e2019430001-01 every four years the geological survey of denmark and greenland (geus) develops and implements new strategies to ensure that we are able to help meet the ever-changing challenges that face society. in 2018 these discussions were shaped by important issues like climate change and climate adaptation, and their consequences for our use of energy, minerals and water resources. as part of this strategic focus, geus introduced a new publication strategy in 2018 that seeks to increase our publication rate of high impact science, and to gain more visibility within the international scientific community and the media. many different tools will be applied to make such a long-term cultural change possible, including modernisation of geus’ own publication series. it is therefore very promising to see this issue of review of survey activities (rosa) published in a modernised bulletin. the geus bulletin is an international open-access publication with peer-reviewed papers. for the first time rosa is published online only and all papers can be downloaded shortly after completion. papers are easily identified and cross-referenced with dois (digital object identifiers), and keywords and publication history are provided. in addition, there are better possibilities for provision of online supplementary data to document large datasets. this new format will hopefully allow for a much quicker and more efficient distribution and thereby attract higher interest from the media and society. this issue of review of survey activities includes 16 papers covering many different activities in denmark, greenland and beyond. six papers are on denmark, eight on greenland, and two on other themes. due to the new format, several of the papers are slightly longer than in previous years, where the four-page limit sometimes made it difficult to fully document and discuss key results. activities in denmark geus activities and research in denmark cover a wide range of topics within our specific programme areas: data, water, energy, mineral resources, nature and climate. our strong focus on climate change and climate adaptation is reflected in several papers. one very interesting paper on sea-level rise in denmark combines the representative concentration pathway 4.5 scenario (“paris agreement” climate pathway) with local reconstructions of glacio-isostatic rebound curves to calculate future differences in sea level between skagen, copenhagen and esbjerg. another paper presents a statistical discussion on how models of precipitation and temperature projections can be downscaled from coarse meteorological grids to the much more detailed grids that geus is using in our danish national water resources model. understanding properties of reservoir sandstones and their fluids is very import for predicting production models of petroleum and geothermal energy, or for carbon capture and storage (ccs). one paper provides an analysis of porosity-preserving microquartz coatings in deeply buried jurassic sandstones in the danish central graben and the implications for production parameters. another paper gives a first classification of different brine types in mesozoic reservoirs in areas with a geothermal potential onshore denmark. this is very important for evaluating the scaling risk prior to new drilling. after many years of unsuccessful onshore petroleum exploration in denmark and the expectation of only limited potential in this area, the danish government finally decided in february 2018 to halt all further petroleum exploration onand nearshore. one paper describes the exploration history in the region since the last major study was completed in 1987, and reviews the various exploration play types. much of the data and results from this decades-long period of onshore exploration are very important for developing a more detailed understanding of the potential for geothermal energy and ccs in many parts of denmark. the last danish paper provides new information on the deeply incised and partly sediment-filled channels in storebælt using radiocarbon ages, and thereby suggests new models for the history of water connections between kattegat and the baltic sea. review of survey activities 2018 flemming g. christiansen*1 deputy director, geological survey of denmark and greenland (geus) geus bulletin is free to individuals and institutions in electronic form. the author(s) retain copyright over the article contents. preface | open access geus bulletin vol 43 | e2019430001 | published online: 07 august 2019 https://doi.org/10.34194/geusb-201943-00-01 https://doi.org/10.34194/geusb-201943https://doi.org/10.34194/geusb-201943-01-01 https://doi.org/10.34194/geusb-201943-01-02 https://doi.org/10.34194/geusb-201943-01-03 https://doi.org/10.34194/geusb-201943-01-04 https://doi.org/10.34194/geusb-201943-01-05 https://doi.org/10.34194/geusb-201943-01-06 https://doi.org/10.34194/geusb-201943-00-01 e2019430001-02 activities in greenland once again, there was a high level of activity in greenland in 2018. geus continues to focus on climate projects related to the promice monitoring programme. furthermore, there are many traditional studies on petroleum and mineral resources, and some new initiatives on geohazards. two papers are on promice services. the first provides a complete mass balance for the greenland ice sheet between 1995 and 2015. data were obtained by repeated measurements of altitude and ice thickness circumscribing the ice sheet at 1708 m altitude – a 5415 km long perimeter survey. the second paper provides an update of the annual calving front lines for 47 marine outlet glaciers distributed throughout greenland. the net area change of these glaciers is significant over the period 1999–2018 and these data will be important for mass balance calculations. mineral exploration by industry has remained relatively low in recent years, and most petroleum activities have stopped since all supermajors and majors have left greenland. it is, however, important to prepare for a future when prices of the most important commodities eventually rise again. one paper provides interesting results on u-pb dating of titanite in paleogene sandstones in a volcanic terrane in east greenland. this can be very important for constraining sources of sedimentary rocks and can help to date the precipitated titanite. another paper is on the liverpool land basement high (llbh), which could offer a large and interesting analogue for fractured basement reservoirs that have recently attracted interest in petroleum exploration in several places in the north atlantic. llbh is well exposed and provides good possibilities for developing 3d models of faults, fractures and veins and their connectivity. sampling may provide information on their origin, e.g. timing and fluid conditions. spaceborne remote sensing data can be a fast and costeffective tool in early mineral exploration. a case study from wollaston forland, north-east greenland is presented. they compare data from aster and sentinel-2 to discriminate between various lithological units. aster seems to give the best results for this mapping purpose. glacial rock flour has recently been suggested as a natural fertiliser that could be used especially in tropical areas with poor soil quality. one paper describes mapping and sampling of such deposits in lake tasersuaq, west greenland. on 17 june 2017 a large landslide in karrat fjord triggered a devastating tsunami. in the weeks that followed, geus was deeply involved in coordinating information and knowledge from many dedicated scientists in order to advise the greenland authorities. it soon became evident that there was a strong need for a better understanding of the risk of landslide-generated tsunamis across greenland. it is particularly important to understand how an apparent increase in the number of events relates to climate change, and how geology, terrain and local climate control the distribution of risk area. in 2018, geus completed a screening study of the risk of major landslides. one paper gives an overview of historical records and applied mapping and satellite techniques that have been used to map the occurrence of more than 500 landslides in greenland. another paper documents a multidisciplinary case study of a landslide in 2018 in the karrat fjord, using seismological data, geological data, optical satellite images and radar satellite data to describe activity before and during the event. other activities internationally, geus collaborates with a number of research institutes on a wide variety of projects, including broader thematic studies and the development of new techniques and databases. based on our own high-quality analytical techniques, geus has worked with provenance studies for several decades and we have compiled large amounts of zircon ages and other types of provenance data in greenland, the north atlantic and the north sea. these data are useful for researchers and industry, and so geus and the norwegian petroleum directorate (npd) have collaboratively launched an online database, the so-called north atlantic provenance database, to collate these regional data. one of the papers in this issue provides an introduction to a number of the visualisation and statistical tools available in the database, which will be developed further in the coming years. the last paper is on the use of drones equipped with multisensors for geological mapping and mineral exploration. the development of light weight magnetic and hyperspectral tools together with a high quality positioning system made in collaboration with several european partners is described, together with a presentation of preliminary results from two mining areas in finland. how to cite christiansen, f.g. 2019: review of survey activities 2018: preface. geological survey of denmark and greenland bulletin 43, e2019430001. https://doi.org/10.34194/geusb-201943-00-01 *corresponding author: flemming g chrisitansen | e-mail: fgc@geus.dk 1 geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350, copenhagen k, denmark. https://doi.org/10.34194/geusb-201943-02-01 https://doi.org/10.34194/geusb-201943-02-02 https://doi.org/10.34194/geusb-201943-02-03 https://doi.org/10.34194/geusb-201943-02-04 https://doi.org/10.34194/geusb-201943-02-05 https://doi.org/10.34194/geusb-201943-02-06 https://doi.org/10.34194/geusb-201943-02-07 https://doi.org/10.34194/geusb-201943-02-08 https://doi.org/10.34194/geusb-201943-03-01 https://doi.org/10.34194/geusb-201943-03-02 https://doi.org/10.34194/geusb-201943-00-01 mailto:bhm%40geus.dk?subject= geological survey of denmark and greenland bulletin 33, 2015, 49-52 49 a quartz-wolframite-molybdenite vein and scheelite in amphibolite horizons from thrudvang peninsula, skjoldungen, se greenland diogo rosa and thomas ulrich during the south-east greenland mineral endowment task (segment) expedition in 2012, the possible mineral potential of the skjoldungen region was investigated. th e region is part of the archaean north atlantic craton, and includes the skjoldungen alkaline province (nielsen & rosing 1990; blichert-toft et al. 1995; kolb et al. 2013). a quartz-wolframite-molybdenite vein with phyllic alteration was identifi ed during the reconnaissance work in the northwestern part of the peninsula of th rudvang, close to the kangertikajik fj ord (ggu 446946; figs 1, 2). th e c. 30 cm wide, subvertical vein is hosted in mafi c granulite. however, the deformed nature of the vein and steep terrain did not allow us to establish its extent or general trend. two rusty amphibolite horizons, in the mafi c granulite hosting the vein, have anomalous concentrations of tungsten, but not of molybdenum (ggu 446934 and 446948; table 1). ultraviolet light examination conducted during the study revealed that tungsten is present as pale blue luminescing scheelite grains, rather than as wolframite, as in the vein. as such, these anomalous amphibolite samples are similar to scheelite-rich stratabound horizons documented in supracrustal sequences in the godthåbsfj ord region in southern west greenland, which were interpreted by appel & garde (1987) to be of exhalative origin. fluid inclusion study a sample from the quartz vein (ggu 446946) was investigated for fl uid inclusions in a thick section. th e quartz is very transparent and shows only very few fl uid inclusions which are generally aligned in secondary trails that are typically parallel and rarely crosscut each other (fig. 3a). th e inclusions vary in size from 5 to 20 mm and are euhedral in shape. only two inclusion trails showed irregular or deformed inclusion shapes. overall, only one type of inclusion was found, characterised by a dark single phase that fi lls the entire inclusion (fig. 3b). selected trails of 24 inclusions from three fl uid inclusion assemblages were used for microthermometry, to determine the phase transition during their cooling. th e dark phase in the inclusions contracted at around –95° to –110°c when a small vapour bubble appeared (fig. 3c). th e inclusions © 2015 geus. geological survey of denmark and greenland bulletin 33, 49–52. open access: www.geus.dk/publications/bull greenland mafic granulite, minor paragneiss, meta-peridotite, amphibolite grey tonalitic to granodioritic gneiss (2781 ± 6 ma in skjoldungen area) tonalitic to granodioritic gneiss, locally agmatitic 1 km 63°30´n 41°w 446934 446946, 446948 t h r u d v a n g dragsfjord kangertikajik jættefjorden fig. 1. geological map of the peninsula of thrudvang and the surrounding areas (after kolb et al. 2013), showing the locations of the samples discussed in this paper. contour line spacing: 100 m. 5050 were then slowly heated at a rate of 3–5°c/min. th e vapour (liquid-gas) bubble homogenised between –59° and –50.4°c forming a single phase. th e range of the diff erent fl uid inclusion assemblages in individual trails is small (fig. 4). th ere was no melting of ice or clathrate observed. th e behaviour of the fl uid inclusions during the microthermometric experiments is typical for high-density liquid fig. 2. top: sampling the vein described in this paper. mid and bottom: field work in the kangertikajik fjord region, south-east greenland in 2012. photographs (mid and bottom): lars lund sørensen. 51 co2 inclusions that homogenise close to the co2 triple point at –56.6°c. th e density of the inclusions is calculated to c. 1.1 g/cm3. th e range of the homogenisation temperatures indicates that a small amount of other elements such as nitrogen is probably present in the fl uid. th e occurrence of one-phase liquid co2 inclusions in metamorphic rocks was interpreted by hollister (1990) and johnson & hollister (1995) to be related to grain boundary migration during re-crystallisation of quartz. th ey proposed that selective removal of h2o from h2o–co2 fl uid mixtures leads to enigmatic pure co2 inclusions. such processes cannot be excluded for ggu sample 446946, but are diffi cult to recognise. th erefore, it cannot conclusively be argued that the co2-rich fl uid is directly related to the w-mo mineralisation observed in these rocks, because it could well be that the fl uid inclusions were modifi ed during the metamorphic history of the sample. geochronology a molybdenite concentrate from the sampled vein was re– os dated and provided a neoarchaean age of 2749 ± 11 ma (table 2). th is age is similar to the laser ablation inductively coupled plasma mass spectrometry (la-icp-ms) date of c. 2740 ma for a porphyritic monzogranite and a laicp-ms date of 2753 ± 5 ma for the skirner bjerge syenite in the skjoldungen alkaline province, which are interpreted to have been emplaced during the fi rst (ds1) stage of regional transpression of the skjoldungen orogeny (kolb et al. 2013). th e vein documents a previously unknown tungstenmolybdenum mineralisation occurrence in greenland. previously known molybdenite occurrences in east greenland are related to palaeogene intrusions. furthermore, the age determination fi ts into the earliest of fi ve mo mineralising pulses, which, according to golden et al. (2013), correspond to supercontinent assembly events. in the case of the neoarchaean mo mineralising pulse, it can be linked to the assembly of kenorland during the neoarchaean, also known as superia (golden et al. 2013). mineral potential th e analysed molybdenite has a relatively low re concentration (table 2), which is typical of archaean molybdenite. th is low re concentration probably refl ects the limited mobility of re in the reducing environment that prevailed prior to oxidation of the atmosphere (golden et al. 2013). notwithstanding the reported secular variation of re concentrations 446934 rusty zone in 49.75 14.61 11.18 0.14 7.39 10.50 3.14 0.65 0.84 0.02 0.89 99.12 3 0.37 amphibolite 446946 w-mo quartz 94.95 1.57 1.10 0.01 0.11 0.36 0.44 0.11 0.05 0.02 0.49 99.2 263 1.60 vein 446948 rusty zone in 49.03 13.99 13.05 0.16 4.77 12.15 2.81 0.30 0.74 0.23 2.18 99.4 5 0.55 amphibolite table 1. whole rock geochemistry of w–mo mineralised samples from thrudvang, se greenland * ggu description sio 2 al 2 o 3 fe 2 o 3 mno mgo cao na 2 o k 2 o tio 2 p 2 o 5 loi total mo w no. % % % % % % % % % % % % ppm % * the samples were analysed at actlabs (canada), using fusion icp-oes (majors) and fusion icp-ms (mo and w). cb a 50 μm 50 μm 200 μm fig. 3. photomicrographs. a: parallel fluid inclusion trails. b: fluid inclusion trails with one-phase liquid co2 inclusions at room temperature. c: the same fluid inclusion trails at –120°c. 5252 in molybdenite, according to stein (2006), the re concentration can be used to establish the type of mineral occurrence and the economic potential. however, the re concentration in the analysed sample is intermediate between that of likely subeconomic molybdenite occurrences formed by local dehydration melting of biotite gneiss (with <20 ppm re or even sub-ppm re), and that of molybdenite of possible economic interest with a porphyry-style intrusion-related origin (with hundreds to thousands ppm re). th erefore, we cannot discriminate between the two mineralisation types and we cannot assess the economic potential of this occurrence, using this criterion. as such, the possibility that the studied vein may be linked to an intrusion and could be part of a wider mineralising system with economic potential remains. finally, bearing in mind the close spatial relations, it is considered that the mineralisation in the rusty amphibolite horizons is contemporaneous with the dated vein, and not of exhalative or syn-genetic origin. in this case, the scheelite in the amphibolite horizons was precipitated in the previously carbonatised mafi c to ultramafi c horizons in the host package, due to their enhanced reactivity to vein-derived mineralising fl uids. th is type of reaction can yield skarn-like occurrences, but they are probably not of economic interest – in contrast to the vein mineralisation, which could be of economic interest. acknowledgements th e work was carried out as part of the segment project, jointly fi nanced by the ministry of industry and mineral resources of the government of greenland and the geological survey of denmark and greenland. references appel, p.w.u. & garde, a.a. 1987: stratabound scheelite and stratiform tourmalinites in the archaean malene supracrustal rocks, southern west greenland. bulletin grønlands geologiske undersøgelse 156, 26 pp. blichert-toft , j., rosing, m.t., lesher, c.e. & chauvel, c. 1995: geochemical constraints on the origin of the late archean skjoldungen alkaline igneous province, se greenland. journal of petrology 36, 515–561. golden, j., mcmillan, m., downs, r.t., hystad, g., goldstein, i., stein, h.j., zimmerman, a., sverjensky, d.a., armstrong, j.t. & hazen, r.m. 2013: rhenium variations in molybdenite (mos2): evidence for progressive subsurface oxidation. earth and planetary science letters 366, 1–5. hollister, l.s. 1990: enrichment of co2 in fl uid inclusions in quartz by removal of h2o during crystal-plastic deformation. journal of structural geology 12, 895–901. johnson, e.l. & hollister, l.s. 1995: syndeformational fl uid trapping in quartz; determining the pressure-temperature conditions of deformation from fl uid inclusions in the formation of pure co2 fl uid inclusions during grain-boundary migration. journal of metamorphic geology 13, 239–249. kolb, j., th rane, k. & bagas, l. 2013: field relationship of high-grade neoto mesoarchaean rocks of south-east greenland: tectonometamorphic and magmatic evolution. gondwana research 23, 471–492. nielsen, t.f.d. & rosing, m.t. 1990: th e archaean skjoldungen alkaline province, south-east greenland. rapport grønlands geologiske undersøgelse 148, 93–100. stein, h.j. 2006: low-rhenium molybdenite by metamorphism in northern sweden: recognition, genesis, and global implications. lithos 87, 300–327. authors’ addresses d.r., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: dro@geus.dk t.u., department of geoscience, aarhus university, høegh-guldbergs gade 2, dk-8000 aarhus c, denmark. 446946 54.02 0.14 33.95 0.09 1591 1 2749 11 table 2. re–os dating of a molybdenite occurrence at thrudvang, se greenland * ggu re 187re 187os model age ± 2σ no. ppm ± 2σ ppm ± 2σ ppb ± 2σ ma ma * carried out at als minerals (canada), using isotope dilution mass spectrometry with a carius-tube, solvent extraction, anion chromatography and negative thermal ionisation mass spectrometry techniques. 5 4 3 2 1 0 fr eq u en cy temperature (°c) –60 –58 –56 –54 –52 –50 fig. 4. histogram of fluid inclusion homogenisation temperatures for three-phase fluid inclusions. geological survey of denmark and greenland bulletin 35, 2016, 13-16 13© 2016 geus. geological survey of denmark and greenland bulletin 35, 13–16. open access: www.geus.dk/publications/bull the subsurface material in urban areas comprises the original geological succession together with anthropogenic modifications and deposits. the geological survey of denmark previously performed geological mapping in selected danish cities (e.g. mertz 1974), but this practice stopped in the mid-1980s. the lack of recent systematic mapping in urban areas is apparent not only in denmark but also in most other european countries (cost 2015). however, there is a growing demand for knowledge of the subsurface beneath our cities for a number of reasons: increased urbanisation, infiltration of excess surface water and other climate-change related measures, thermal storage, groundwater cooling and abstraction, subsurface infrastructure, infrastructure projects, etc. the physical properties of the subsurface material are in constant change due to urban growth and infrastructure development. this can strongly influence the geotechnical properties and handling of excess surface water. in order to manage both challenges and opportunities of the ground beneath the cities there is a growing need for 3d hydrogeological models that can encompass all relevant parts of the physical subsurface system and act as operational tools in its management. with the main focus on hydrogeology and the urban water cycle, the municipality of odense, the local waterworks (vandcenter syd), the geological survey of denmark and greenland (geus) and two consultants (alectia and i-gis) have made a joint effort to systematically map the subsurface layers and build a 3d hydrogeological model of the subsurface of the city of odense (fig. 1). this paper provides an overview of the project rationale and an outline of the major results. the sedimentary succession beneath odense the uppermost 50 to 100 m of the subsurface of odense is dominated by weichselian clay till with intervening sand layers (e.g. jørgensen & piotrowski 2003, mertz 1974), which form three groundwater reservoirs of varying extent and thickness. there are also lateto postglacial, nearsurface, sandy outwash plains and heterogeneous infill of erosive channels and depressions (sandersen et al. 2015). historical maps from the late 1800s show that postglacial bogs and wetland areas have also earlier been present. the urban development of odense has mainly taken place within the past 200 years (fig. 1). the switch to the industrial era increased its population and led to expanding residential and paved areas, installation of water supply and sewage systems, creation of waste dumps and an accelerated abstraction of groundwater. the drainage and lowering of the groundwater table dried out some of the former wetlands and created new agricultural and urbanised land, and in other areas caused subsidence due to the decay of organic matter. since the 1980s, a growing environmental awareness and increased taxes have resulted in a marked decrease in groundwater abstraction, and former 3d hydrogeological modelling for urban subsurface management in odense, denmark susie mielby, tom martlev pallesen and peter b.e. sandersen 1717 1839 1892 1928 1977 urban plan 2009–2021 odense municipality odense od ens e fj ord 5 km fig. 1. urban development of odense municipality (laursen & mielby 2016). 1414 wetland areas are beginning to return to their original wet state. the urban activities within the last 200 years created a heterogeneous man-made layer consisting of backfill of excavations, modified terrain, landfill, waste dumps, etc. parts of this anthropogenic layer rest on top of the original geological formations, whereas other parts replace them. building the 3d model the man-made modifications of the urban subsurface, along with the ongoing climate change, affect the urban water cycle and must therefore be carefully evaluated. for instance, in areas with sustainable urban drainage systems (suds), the local infiltration potential and its consequences on the hydraulic heads in adjacent areas must be known (jeppesen 2014; mielby et al. 2015a), and this implies a need for a close integration of regional and local geological information. in order to build a comprehensive subsurface model for odense both the natural and man-made layers were mapped, but due to their different nature, different mapping and modelling approaches were needed. a standard framework-model approach where layer boundaries were mapped all the way to the surface was employed for the original geology, primarily using borehole data, geophysical data and high-resolution digital elevation models (sandersen et al. 2015). the man-made layers shown in fig. 2 were mapped using data from boreholes about the thickness and character of infill, and from excavations around subsurface infrastructures. the latter requires detailed information about the age, character and spatial extent of the individual parts of the infrastructure and access to current standards for excavation back-fill (fig. 3; pallesen & jensen 2015). the digital information about the anthropogenic layer is generally abundant, but also rapidly changing compared with the traditional geological data. because the elements of the anthropogenic layers change over time, a tool for the modelling of the anthropogenic layers was established in the modelling software geoscene3d to make a sequential handling of the infrastructure data possible (pallesen & jensen 2015). the need for a high degree of detail is most relevant in the uppermost parts of the 3d model, where the scale of the urban infrastructure typically is counted in metres or tens of metres. mapping of the man-made layers should therefore be done with a corresponding level of detail. in the deeper parts of the subsurface, the number of boreholes that reach depths of 20 m or more is very limited (kristensen et al. 2015), meaning that the deeper geological succession typically cannot be resolved better than 100 m horizontally (sandersen et al. 2015). after mapping of the original geological layers and the man-made components the two parts are merged, whereby the latter component replaces the model part of the original geological layers (fig. 4). in this way, two models are combined into one 3d model. however, due to the above-mentioned scale difference, it is important to choose a common resolution that meets the required need for detail but does not exceed computational capability. the major part of the geological information in the urban area comes from relatively short boreholes drilled in connection with geotechnical and environmental projects. the geotechnical borehole information is very important, as it provides information about both the original geology and the composition of the man-made fill. data from many of these boreholes are stored in the national jupiter database, but its quality is variable because upload of this type of borehole data is not mandatory. in addition, many boreholes have not found their way to the national database at all, because the data were collected by private companies, considering the information their private property. therefore, the readily available digital geological information does not necessarily reflect all of the existing data, and this problem must be addressed before a 3d model is constructed. fig. 2. simplified picture of the subsurface elements in the man-made urban layer, where the original geological sediments (light grey) are replaced by infill (yellow), basements of buildings, wells, pipes, sewers, etc. forming the anthropogenic layer. 15 results the project resulted in a 3d geological municipality model, and detailed anthropogenic and combined hydrogeological models were made for a chosen site in order to test the modelling in a typical situation. in addition to the models and the model tool in geoscene3d, descriptions, workflows and recommendations for data acquisition, management and updating procedures were also developed. in the project, a systematic and comprehensive collation of available data was established within the borders of odense municipality. in 2015, the geological model concept provided the ‘best so far‘ foundation for the management of the urban hydrogeology in odense. the resulting model and its tools were proven to be valuable, as they were quickly taken into use in several new hydrological modelling projects such as modelling of hydrological climatechange effects and detailed infiltration possibilities, as well as the evaluation of the natural protection of drinkingwater well fields. the data and information density proved variable, and therefore information on the data background is important for the municipal decision-making. detailed studies in odense showed that in many areas the data coverage is insufficient for the purpose required. therefore, the geological basis for planning and water management has to be improved, and this requires a dynamic model that can benefit from both existing and new data. if all additional geological data are to be accessible for future planning, the authorities must ensure that all geological and geotechnical data are reported and available. the investigations in the current project proved that open access to better geotechnical and geological data and modelling would also benefit a wider range of users such as archaeologists, engineers, architects, entrepreneurs and other professionals working with e.g. road and railway construction (laursen et al. 2015). if the 3d model is regularly updated in the future, odense will also obtain an increasingly robust foundation for the hydrological modelling and management of the urban water cycle. detailed recommendations and experience from the current project (mielby et al. 2015b), urban data modelling tools in geoscene3d and data storage facilities for geotechnical information in the jupiter database (hansen et al. 2015) are available for the benefit of other areas. a b fig. 3. a: large sewer in an excavation to be filled with gravel. image source: vandcenter syd. b: anthropogenic model of sewers with other implemented data elements (roads, pipes and buildings). 1616 acknowledgements the foundation for development of technology in the danish water sector (vtufonden; ans-7497.2012) is thanked for financial support. the project partners knud søndergaard and gert laursen, odense municipality, christian ammitsøe and johan linderberg, vcs denmark, martin hansen and margrethe kristensen, geus and jan jeppesen, alectia are acknowledged for fruitful collaboration. references cost 2015: http://www.cost.eu/cost_actions/tud/actions/tu1206. suburban – a european network to improve understanding and use of the ground beneath our cities. hansen, m., wiese, m.b., gausby, m. & mielby, s. 2015: udvikling af en 3d geologisk/hydrogeologisk model som basis for det urbane vandkredsløb. delrapport 6 – teknisk håndtering og lagring af bygeologiske data og modeller, 22 pp. københavn: de nationale geologiske undersøgelser for danmark og grønland. jeppesen, j. 2014: udvikling af en urban-hydrologisk model til simulering af nye innovative lar-løsninger til lokal håndtering af både regn-vand og grundvand (larg). afrapportering af vtu-projekt 29. december 2014. jørgensen, f. & piotrowski, j.a. 2003: signature of the baltic ice stream on funen island, denmark during the weichselian glaciation. boreas 32, 242–255. kristensen, m., sandersen, p. & mielby, s. 2015: udvikling af en 3d geologisk/hydrogeologisk model som basis for det urbane vandkredsløb. delrapport 2 – indsamling og vurdering af data, 82 pp. københavn: de nationale geologiske undersøgelser for danmark og grønland. laursen, g. & mielby, s. 2016: odense. tu1206 cost sub-urban wg1 report. laursen, g., mielby, s. & kristensen, m. 2015: udvikling af en 3d geologisk/hydrogeologisk model som basis for det urbane vandkredsløb. delrapport 3 – geotekniske data til planlægning og administration, 32 pp. københavn: de nationale geologiske undersøgelser for danmark og grønland. mertz, e.l. 1974: odense og omegns jordbundsforhold: en ingeniørgeologisk beskrivelse. danmarks geologiske undersøgelse rapport 9, 37 pp. mielby, s., laursen, g., linderberg, j., sandersen, p. & jeppesen, j. 2015a: udvikling af en 3d geologisk/hydrogeologisk model som basis for det urbane vandkredsløb. delrapport 1 – 3d-modellen som basis for håndteringen af det urbane vandkredsløb, 66 pp. københavn: de nationale geologiske undersøgelser for danmark og grønland. mielby, s., jespersen, c.e., ammitsøe, c., laursen, g., jeppesen, j., linderberg, j., søndergaard, k., kristensen, k., hansen, m., jensen, n.-p. & sandersen, p. 2015b: udvikling af en 3d geologisk/hydrogeologisk model som basis for det urbane vandkredsløb. synteserapport, 57 pp. københavn: de nationale geologiske undersøgelser for danmark og grønland. pallesen, t.m. & jensen, n.-p. 2015: udvikling af en 3d geologisk/ hydrogeologisk model som basis for det urbane vandkredsløb. delrapport 5 – interaktiv modellering af antropogene lag, 58 pp. københavn: de nationale geologiske undersøgelser for danmark og grønland. sandersen, p., kristensen, m. & mielby, s. 2015: udvikling af en 3d geologisk/hydrogeologisk model som basis for det urbane vandkredsløb. delrapport 4 – 3d geologisk/hydrostratigrafisk modellering i odense, 106 pp. københavn: de nationale geologiske undersøgelser for danmark og grønland. a b c fig. 4. elements of the geological modelling. a: buildings and pipes. b: man-made ground with fill. c: combined model showing man-made features and the underlying geological model. authors’ addresses s.m. & p.s., geological survey of denmark and greenland, c.f. møllersvej 8, building 1110, dk-8000 aarhus c, denmark. e-mail: smi@geus.dk. t.m.p., i-gis, voldbjergvej 14, dk-8240 risskov, denmark. geological survey of denmark and greenland bulletin 38, 2017, 69-72 69 denmark has a long tradition for having central geological databases, including a systematic collection and storage of geological and hydrological information from all surficial boreholes which was initiated in 1926. since the mid-1970s such data have been stored digitally. a large variety of users access a central danish, geological database: the public, for information about their local drinking water quality, environmental employees in municipalities, regions and the state for using, entering and updating data as well as consultants and drilling companies working for public administration and local water works. the local danish administrative system previously consisted of 14 counties and 248 municipalities. the counties were responsible for groundwater mapping, drinking water management and activities concerning contaminated soil, as well as for harmonisation and transfer of data to the central database. with effect from 1 january 2007, this administrative system was replaced by five regions, seven environmental centres and 98 municipalities, which required major changes in the administrative handling of borehole data at the local and regional levels. for this, a public and shared central database was established and a countrywide harmonisation of data, transfer and storage was initiated and all geological, groundwater and drinking water data were transferred to this central database at geological survey of denmark and greenland (geus). in an updated database system, public authorities were set up to access the central database to store their relevant borehole data and almost all data were made publicly available. the database is maintained by geus. it is directly connected to other public databases at geus including the shallow geophysical database gerda, where e.g. borehole loggings are stored, and to the model database where simple geological models are stored (fig. 1). an integrated public information system for geology, groundwater and drinking water in denmark martin hansen and charlotte toftemann thomsen fig. 1. diagram showing a: the access to data in the jupiter database from different users and b: how data flow to and from the database. ws: web services. privileged users jupiter database pc jupiter xl read-only views jupiter tables unprivileged users users at geus geus databases drilling companies municipalities state agencies consultants borehole data groundwater data borehole data models reports borehole data geophysics models reports groundwater and drinking water data laboratories chemical analysis regions b. data flow a. access to jupiter data http – homepages, ws reading and download web forms updating using ws and https borehole, © 2017 geus. geological survey of denmark and greenland bulletin 38, 69–72. open access: www.geus.dk/publications/bull 7070 this updated system is used by the municipalities to manage their water supply data (e.g. water supply structure, permits, groundwater and drinking water quality data), by state agencies to manage groundwater data from the groundwater mapping and by the regions to maintain their soil pollution data. the system has gradually been expanded since 2007 and now local authorities can store and maintain a wide range of their own data in the central database. the database system the database system currently consists of the following components: 1. a central database; 2. a public data model, agreed upon by a committee under the the danish natural environment portal; 3. a user-management system, managed by the the danish natural environment portal, providing direct access to the central database; 4. a suite of simple object access protocol (soap) web services – an interface that allows computer-to-computer communication. this enables local authorities to manage their own data in the central database through their own applications; 5. applications that can access the database utilising the components above. the central database the central database is based on geus’ jupiter database and run on an oracle database. it has been under development during the last 40 years. the extensions made since 2007 include full public online access to read data and write access for public authorities to almost all data types. the public part of the database is made available through a view layer exposing the public data model. the public data model the public data model can handle: • borehole data, including: localisation and administrative data, construction data, abandon data, geological description using common methodology, hydraulic head measurements, samples and analyses from soil, water and air; • surface soil and water sample data; • water supply data for water plants, agriculture and industry, including: extractions wells, well fields, water treatment plants, extraction permits, water quality data, water extraction, water use data, exchange of water between water works, ownership and contact persons; • soil pollution data: projects; soil, water and air chemistry data from boreholes, surface samples and remediation plants. the data model is being developed and maintained by geus, but all extensions and alterations have to be agreed upon by the groundwater group under the danish natural environment portal (dnep). this is a common public partnership between the ministry of environment and food of denmark (45%), the danish municipalities (45%) and the danish regions (10%), and it acts as an independent portal across boundaries of authority. its major goal is to ensure continued access to harmonised, updated, natural environmental data. the groundwater group itself consists of members appointed by the municipalities, name of user role privileges and use laboratory for entering and editing samples and chemical analyses. for laboratories to add data to the system. borehole write for entering and editing information about boreholes including location, geology/lithology and borehole construction. for consultants and local authorities to enter new boreholes. borehole read for reading publicly available data about bore holes, water, soil and air samples and analysis. drinking water for approving new drinking water samples with approval analyses. for users from the authorities to qua lity control new drinking water data and by approving the data, making them publicly available. groundwater for approving new groundwater samples with approval analysis. for users from the authorities to quality control the new groundwater data and by approving the data, making them publicly available. water level for entering and editing water-level measurements. for users from the authorities to enter and edit water-level measurements. sample approval for approving new water, soil and air samples with analyses from surface samples. for users from the authorities to quality control the new water, soil and air analyses from surface samples and by approving the data, making them publicly available. water resources for entering and updating information about water works. these data include water well fields, treatment plants, permits, annual volumes of extracted water, sampling sites and volumes of water shared between different water works. for the municipalities that survey the drinking water to create and update their drinking water structure. table 1. user roles and privileges 71 the danish environmental protection agency , the danish regions, dnep and geus. the public data model consists of more than 90 data tables. data responsibility agreement the data responsibility agreement determines which organisations are responsible for producing and maintaining which data, and for making the data available to the public. the various responsibilities are defined partly by legislation and partly through agreements signed by the participating partners and by voluntary reporting. at any time, any data set in the public database has one and only one responsible owner organisation. data ownership the ownership of data can be defined either by the user who enters the data or by the location of the data point. in this way, the municipality or region in question, a state agency and geus can own water-level measurements in the same well. each data owner is responsible for entering their own data and secure their quality. the analysing laboratory carrying out the quality control of drinking water is responsible for entering its data, and it owns the data until the data entry and quality control have been completed. after this step, the ownership is transferred to the municipality to which the water works belongs. the municipality has to release the data before it becomes publicly available. apart from quality control of the data sets, the new owner cannot alter the data. if errors are found during the quality control, the municipality must reject the water sample and all of its analyses, and the laboratory has to resubmit a corrected data set. user management the danish natural environment portal has a central user-management system which enables the user to use the same login credentials to access and update data in different systems. the users and their rights are managed locally by user administrators who define the user rights through a set of roles. each role defines to which part of the database the user shall have access. in this way it is the local administrator who decides who should be allowed to access the different public systems or obtain privileges to enter and edit data. the different roles in the jupiter system are presented in table 1. in addition to these roles, the system gives the users access to data according to their geographical location. for example, all users can access water works but only users from the municipality, where the water works is situated, can update the information (provided the user has the right role). data interface the data are available in several ways. most of the data are available through: jupiter’s homepage: this entry is read-only and mainly used by municipalities and members of the public to look up specific data (fig. 2). soap web services: these services give read-and-write access to different parts of the data model and have been under development since 2007. with this interface private companies can write applications for administrative units for their data management. these services constitute fig. 2. screen shot from the web map interfaces to jupiter showing boreholes and water plants from central sjælland. 7272 the main entry outside geus for update of data. see for example: http://webs.geus.dk/miljoeportal.groundwater.bboring.2.0.0/b-boring?wsdl for updating borehole data. wms/wfs: several of the data themes are exposed through web map services (wms) and web feature services (wfs). web map services deliver maps as bitmaps while web feature services can deliver the same data as geographical objects (point, lines and polygons) that can be used for spatial queries in a gis. these are used to support the map interface on the jupiter homepage, are available for end users, and can also be imported into local gis. these services are used mainly in systems made for the different administrating units. see for example the wms publishing borehole information http://data.geus.dk/geusmap/ows/25832.jsp?se rvice=wms&version=1.1.1&request=getcapabil ities&layers=jupiter_boringer_ws%2cjupiter_anlaeg_ws database download: advanced users can download data as database exports. in this way it is possible to export all available data to a local hard drive (excluding water quality control data not yet approved by the data owner and information about owners and contact persons). this function is meant for advanced users for e.g. geological modelling or complex calculations on groundwater chemistry. it is even possible to install a scheduled application that keeps the local database updated on a nightly basis with changes made in the central database. such local copies of the database are mainly used by consulting companies and large administrative units. discussion and conclusions nearly all data in the database must be publicly available. therefore, the access has been divided into two packages of services since the first version of the web services was developed up to january 2007. one set contains all the readonly functionality without any user management systems, while the other package contains functions for updating the data. however, due to very frequent use the read-only services will have to be revised in the near future. not all users comply with the rules set up for the use of the services, and since they are anonymous, it is difficult to identify those who break the rules. for example, users are not allowed to use the services in batch mode or create a local copy of the database. we can, however, see from the logs, that one or more read-only users behind a single ip-address make up to tens of thousands of calls on a daily basis and thus obviously do not comply with the rules set-up for the services. an increase from c. 200 000 to 12 000 000 calls per month in the last few years causes a heavy and increasing system load. if a login with user name and password was to be required to enter the read-only services, it would be possible to contact directly the users who use software that does not comply with the rules of use. the access would still be free of charge. a public, shared database like jupiter gives access to a very broad use, where the data can be combined with other public data or with private, non-public data. also the many different ways in which the data are available, such as web, web gis, different types of web services or download in database format, make the data highly usable. the user gets a coherent dataset containing geology, groundwater and drinking water data, where the water can be followed all the way from the borehole to the water plant. in recent years, the database has been used for analysis of public health in combination with drinking water quality. the free access to the publicly available data has greatly increased the value of the data. the authors do not know of any other publicly available, combined geology – groundwater – drinking water database systems like jupiter. as the system is based on a data model that has been agreed upon between different stakeholders from municipalities, regions, state agencies and the geological survey, the model can most probably be used as a good starting point for development of similar systems by other organisations and countries. references gerda database: http://data.geus.dk/geusmap/?lang=en&mapname=gerd a#layers=gerda_ projects%2cgerda_data jupiter database: http://data.geus.dk/geusmap/?mapname=jupiter&lang=en saml 2.0: https://en.wikipedia.org/wiki/saml_2.0 soap web services: https://en.wikipedia.org/wiki/soap the danish natural environment portal: http://www.miljoeportal.dk/ english/sider/default.aspx the model database: http://data.geus.dk/geusmap/?lang=en&mapname =modeldb authors’ address: geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: mh@geus.dk geological survey of denmark and greenland bulletin 33, 2015, 73-76 73© 2015 geus. geological survey of denmark and greenland bulletin 33, 73–76. open access: www.geus.dk/publications/bull digital models based on images taken with handheld cameras – examples on land, from the sea and on ice erik vest sørensen, morten bjerager and michele citterio geological outcrops can be comfortably modelled in three dimensions in the offi ce using images from a handheld digital camera. recent developments within the imaging techniques of structure from motion (lowe 2004; snavely et al. 2008; fonstad et al. 2013) and photogrammetry (hirschmüller 2005; james & robson 2012; favalli et al. 2012) have made it easier and cheaper to construct so-called digital outcrop models using stereoscopic images from standard digital cameras. th e digital outcrop model (bellian et al. 2005) is a 3d representation of the outcrop surface and is oft en displayed in the form of a polygon mesh or a point cloud. in this paper we present three examples of such point clouds from images obtained with a handheld digital camera. th e examples illustrate how outcrop topography or digital outcrop models can be constructed at diff erent scales, with diff erent accessibility and operational platforms. two examples illustrate outcrop scales of metres to kilometres, with images obtained by walking along excavated exposures in the faxe limestone quarry and from a boat sailing past the coastal cliff of stevns klint. th e third example illustrates detailed micro-topography of ice and snow surfaces where the images were obtained from a snowmobile on an ice cap in a.p. olsen land, north-east greenland. methods th e images were collected with a 36 megapixel nikon d800e camera equipped with a fi xed 35 mm f/1.4 zeiss lens. th e camera was locked at infi nity in the faxe quarry and stevns klint examples. in the third example from greenland, the camera was focused and locked so that objects at a distance of c. 2 m were in focus. th e images were recorded with ste55°15´20´´n 55°15´20´´n 12°7´30´´e 200 m 12°7´30´´e fig. 2a fig. 1. overview of the faxe limestone quarry. red dots: images obtained in 2013. green dots: images obtained in 2014. the inset map shows the location of the faxe quarry in denmark. orthophotograph from danish geodata agency. fig. 2. a: perspective view (towards the north-east) of the northern part of the faxe limestone quarry showing the constructed point cloud generated from oblique images obtained in 2014. the point cloud is coloured according to colour value of matched pixels and can be rotated freely in three-dimensions. for location see fig. 1. b: close-up of a bryozoan limestone mound with nodular flint layers. b a fig. 2b 20 m 5 m 7474 reoscopic overlaps of up to 90%. in this way image acquisition for digital outcrop models diff ers from the traditional approach for 3d stereoscopic work (dueholm 1992), where a stereoscopic overlap of 60–80% is suffi cient to ensure good precision and continuous stereoscopic overlap. furthermore, images were acquired from much more varied image positions, for example in the faxe quarry study (fig. 1), compared with traditional 3d mapping, where images are typically collected along straight parallel lines. th e construction of the digital outcrop models from the images is based on automatic dense multi-view, stereomatching routines. th ese routines attempt to match each pixel across a range of images. because of the large overlap, the image baseline is quite small, which decreases the precision of matched pixels. th is is, however, compensated for by the redundancy of determining the same point in multiple images. in practice this approach yields a level of precision which is comparable to that typically obtained with stereoimages with 60% overlap, but with a much better automatic elimination of erroneously matched pixels. th is leads to the production of dense point clouds, which require little manual editing, making them well suited for visualisation. th e clouds can also be used in morphological analyses. a number of soft ware solutions can generate point clouds from images; in this study we used the professional version of agisoft photoscan and sure – photogrammetric surface reconstruction from imagery. walking along exposures – the faxe limestone quarry danian deep-water bryozoan and coral carbonate mounds are exposed in the faxe limestone quarry (fig. 1; 55°15́ 40˝n, 12°07´20˝e), which represents a perfect case study for 3d outcrop modelling. th e quarry was visited in 2013 and 2014, and a large collection of stereoscopic images documents the changing features of the active quarry. images of the quarry walls were collected using a handheld digital camera from a distance of 10–20 m, which translates into images with pixel sizes on the ground, also known as the ground sampling distance, in the millimetre range. we used a subset of the data from the northern part of the quarry that is being actively quarried. th e result of the reconstruction is a dense point cloud (fig. 2), which can be freely rotated in 3d and zoomed in on areas of interest. th is is a powerful way of visualising geological outcrop data. when combined with a periodic recording of digital images in the active quarry it can provide unique outcrop topographic data sets that make a reconstruction of the 3d mound topography possible in great detail. it also provides a data set that can be used to quantify volumes of specifi c characteristic rock types, such as the amount of black-grey nodular fl int in the greyish white bryozoan limestone mound systems. b a 10 m fig. 3. a: perspective view of the point cloud generated from a section with danien bryozoan limestone mounds in the coastal cliff of stevns klint. the point cloud is illustrated with rgb-values of matched pixels. b: perspective view of the filtered point cloud based on colour and surface roughness calculations. the inset map shows the location of stevns klint in denmark. fig. 4. perspective view of the reconstructed snow surface measuring 420 × 160 cm from the ice cap in a.p. olsen land, north-east greenland. footprint ruler 50 cm 75 sailing along exposures – stevns klint stevns klint (55°15́ 39˝n, 12°24́ 52˝e) was recently included in unesco’s world heritage list and is world famous for its excellent exposure of the cretaceous–palaeogene boundary. images were collected along 11 km of the coastal cliff from a boat in june 2014, adding new oblique images to the growing archive from previous studies (surlyk et al. 2006; pedersen & damholt 2012). th e images were collected from a distance of 10–300 m, which translates to pixel sizes in the millimetre to centimetre range. th is approach allows for high resolution mapping of the mound structures and megaand mesoscale bedding. images from a cliff section at stevnsfortet in the southern part of the cliff were selected to illustrate how simple manipulation of the generated point cloud data (fig. 3a) can be used to visualise the overall mound structures. characteristic black and grey fl int nodules follow the internal bedding of the bryozoan mounds and display a strong colour contrast to the light-coloured limestone that can be used to fi lter away light-coloured points. th e resulting point cloud can be analysed with calculations of outcrop parameters such as surface roughness and curvature or, with more sophisticated calculations, used in semi-automatic tracing of discontinuities such as joints, fractures or bedding, developed for terrestrial lidar data (garcia-sellés et al. 2011). th e result of the fi ltering is shown in fig. 3b, which illustrates the structure of the internal bedding in a diff erent way. with little eff ort this can be extended to include the entire 11 km surveyed part of the coastal cliff . standing on a snowmobile on an ice cap – north-east greenland th e third example is from an ice cap in a.p. olsen land (74°37´28˝n, 21°22´30˝w) in north-east greenland. th e small-scale topography of snow and ice infl uences the turbulent and radiative components of the surface energy balance, as it controls the aerodynamic roughness length and changes the surface albedo (munro 1989; warren et al. 1998; brock et al. 2006). surface roughness must also be accounted for in remote sensing of the cryosphere (könig et al. 2001). th e spatial scales relevant for such applications span several orders of magnitude, with required vertical accuracies in the order of millimetres (rees & arnold 2006). ground-based photogrammetry appears to be a viable technique to map snow and ice micro-topography in the fi eld down to a scale of centimetres (irvine-fynn et al. 2014). our aim is to demonstrate the feasibility of millimetre-scale accuracy over an outcrop scale of several square metres, under fi eld conditions. 50 cm a b c d 0.055 0.051 0.048 0.044 0.041 0.038 0.034 0.031 0.027 0.024 0.021 0.017 0.014 0.010 0.007 0.003 0 roughness elevation radius = 0.005 m radius = 0.05 m radius = 0.5 m 653.049 653.041 653.034 653.026 653.019 653.011 653.003 652.996 652.988 652.981 652.973 652.966 652.958 652.951 652.943 652.936 652.928 (m above sea level) fig. 5. point cloud of the surface shown in fig. 4, in plane view. a: coloured according to elevation height. b, c, d: coloured according to surface roughness. the surface roughness is calculated as the vertical deviation of each point from the best fitted plane defined by data within a sphere with a radius of 0.005 m (b), 0.05 m (c) and 0.5 m (d). 7676 for this experiment, a person standing on a snowmobile and pointing the camera obliquely downwards collected 29 images from diff erent positions. th e distance from the camera to the ground was c. 2 m. th is approach gives images with a ground sampling distance in the sub-millimetre range. th e generated data are a very dense point cloud (fig. 4). th e height of roughness elements per unit length is shown in fig. 5 over three diff erent spatial wavelengths of 0.005 m, 0.05 m and 0.5 m. th is case study shows that it is possible to obtain surface roughness data useful for glaciological and remote sensing applications by relatively simple means. summary th is study demonstrates that it is possible to generate highresolution topographic data at various scales with diff erent accessibility and operational platforms by using a standard digital camera and computer soft ware. th e method has a high potential for fi eld geologists, who wish to establish accurate outcrop topographic models that can be ‘brought to life’ and visualised in 3d surface models. th ese models can be freely rotated in three dimensions and are well suited for visualisation as well as quantitative purposes in geological mapping. th is is an important new addition to the way 3d mapping is undertaken in the photogrammetry laboratory at the geological survey of denmark and greenland. acknowledgements data from the faxe limestone quarry and stevns klint were obtained with support from geocenter denmark and the european science foundation cocarde-ern. references bellian, j.a., kerans, c. & jennette, d.c. 2005: digital outcrop models: applications of terrestrial scanning lidar technology in stratigraphic modeling. journal of sedimentary research 75, 166–176. brock, b.w., willis, i.c. & sharp, m.j. 2006: measurement and parameterization of aerodynamic roughness length variations at haut glacier d’arolla, switzerland. journal of glaciology 52, 281–297. dueholm, k.s. 1992: geologic photogrammetry using standard smallframe cameras. rapport grønlands geologiske undersøgelse 156, 7–17. favalli, m., fornaciai, a., isola, i., tarquini, s. & nannipieri, l. 2012: multiview 3d reconstruction in geosciences. computers & geosciences 44, 168–176. fonstad, m.a., dietrich, j.t., courville, b.c., jensen, j.l. & carbonneau, p.e. 2013: topographic structure from motion: a new development in photogrammetric measurement. earth surface processes and landforms 38, 421–430. garcía-sellés, d., falivene, o., arbués, p., gratacos, o., tavani, s. & muñoz, j.a. 2011: supervised identifi cation and reconstruction of nearplanar geological surfaces from terrestrial laser scanning. computers & geosciences 37, 1584–1594. hirschmüller, h. 2005: accurate and effi cient stereo processing by semiglobal matching and mutual information. in: schmid, c., soatto, s. & tomasi, c. (eds): proceedings of ieee conference on computer vision and pattern recognition (cvpr), san diego, ca, usa, 20–26 june 2005, 2, 807–814. irvine-fynn, t.d.l., sanz-ablanedo, e., rutter, n., smith, m.w. & chandler, j.h. 2014: instruments and methods. measuring glacier surface roughness using plot-scale, close-range digital photogrammetry. journal of glaciology 60, 957–969. james, m.r. & robson, s. 2012: straightforward reconstruction of 3d surfaces and topography with a camera: accuracy and geoscience application. journal of geophysical research 117(f3), f03017. könig, m., winther, j.g. & isaksson, e. 2001: measuring snow and glacier ice properties from satellite. reviews of geophysics 39, 1–27. lowe, d.g. 2004: distinctive image features from scale-invariant key points. international journal of computer vision 60, 91–110. munro, s. 1989: surface roughness and bulk heat transfer on a glacier: comparison with eddy correlation. journal of glaciology 35, 343–348. pedersen, s.a.s. & damholt, t. 2012: cliff collapse at stevns klint, south-east denmark. geological survey of denmark and greenland bulletin 26, 33–36. rees, w.g. & arnold, n.s. 2006: scale-dependent roughness of a glacier surface: implications for radar backscatter and aerodynamic roughness modelling. journal of glaciology 52, 214–222. snavely, n., seitz, s. & szeliski, r. 2008: modeling the world from internet photo collections. international journal of computer vision 80, 189–210. surlyk, f., damholt, t. & bjerager, m. 2006: stevns klint, denmark: uppermost maastrichtian chalk, cretaceous–tertiary boundary, and lower danian bryozoan mound complex. bulletin of the geological society of denmark 54, 1–48. warren, s.g., brandt, r.e. & hinton, o.p. 1998: eff ect of surface roughness on bidirectional refl ectance of antarctic snow. journal of geophysical research 103(e11), 25789–25805. authors address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail address: evs@geus.dk geological survey of denmark and greenland bulletin 31, 2014, 63-66 63 aeromagnetic survey in south-eastern greenland: project aeromag 2013 peter riisager and thorkild m. rasmussen aeromagnetic surveys are nowadays used at a wide range of scales and purposes. in frontier and under-explored areas, where data are otherwise sparse or non-existent, aeromagnetic acquisition remains the cheapest and easiest way to obtain or refine a picture of the structural setting. aeromagnetic data are also useful for strategic planning of geological mapping campaigns and detailed geophysical data acquisition. moreover, aeromagnetic data are of importance for prospecting, helping to define prospects. large aeromagnetic surveys can be carried out efficiently and safely almost everywhere, in a short period of time and at reasonable cost. in the following we present the newly released aeromag 2013 aeromagnetic survey that covers a remote and relatively under-explored coastal region in south-eastern greenland, stretching from 64°45́ n and northward to 67°30´n (fig. 1). the survey represents a total of 65 492 line km, and covers an area of 30 100 km2, adding a significant new dataset to the already existing database of government-financed geophysical surveys in greenland. with the completion of the aeromag 2013 project, the database presently contains a total of c. 633 500 line km of high-resolution aeromagnetic data and c. 75 000 line km of multi-parameter data (electromagnetic, magnetic and partly radiometric data). further details on previous surveys on greenland and the database of available aeromagnetic data are summarised in rasmussen et al. (2013; see also fig. 1). details of the aeromag 2013 survey eon geosciences inc. flew the aeromag 2013 survey between 13 june and 2 september 2013, using two piper navajo pa-31 aircraft equipped with geophysical instruments (detailed below), and operating out of the airport at kulusuk. the magnetic base stations used for correction of diurnal magnetic variations were installed at two different locations in kulusuk. the survey was carried out by flying along a gently draped surface 300 m above the ground or sea level. due to the severe topography ranging from sea level to c. 2450 m, the gentle drape resulted in an average height above ground of 711 m. the survey lines were ne–sw-oriented, parallel to the coastline with a separation of 500 m while orthogonal tie-lines were flown with a separation of 5000 m. total magnetic field data were recorded with a sampling interval of 0.1 sec. which corresponds to a sample distance of c. 7 m. the magnetic field at the base station was recorded with a 1 sec. sampling interval. aircraft positional data from differential gps measurements were recorded with a 1 sec. sampling interval, and aircraft altitude measurements obtained from barometric altimeter and radar were recorded with a sampling interval of 0.1 sec. a continuous digital video recording of the terrain passing below was also produced. further details on the survey operation and equipment can be found in a report by eon geosciences inc. (2013), which is available at the online dodex database at the geological survey of denmark and greenland (riisager et al. 2011). © 2014 geus. geological survey of denmark and greenland bulletin 31, 63–67. open access: www.geus.dk/publications/bull aem 1998 aem 1994 aem 1995 aem 1996 aem 1997 aem 1998 aeromag 2001 aeromag 2013 aeromag 2012 aeromag 1997 aeromag 1992 aeromag 1998 aeromag 1996 aeromag 1995 aeromag 1999 80°n 75°n 65°n 60°n 80°w 60°w 40°w 50°w 40°w 20°w 20°w 0°w greenland 500 km 60°n 75°n 65°n 70°n 70°n fig. 1. map of greenland showing the location of government-financed high-resolution airborne geophysical surveys conducted from 1992 to 2014. red: aeromagnetic surveys (aeromag). blue: combined electromagnetic and magnetic surveys (aem). 6464 results and products aeromag 2013 survey dataset can be obtained as line geomagnetic reference field corresponding to the date and location of the measurement has been subtracted from the data leaving the magnetic anomalies caused by the variation of magnetisation in the uppermost crust. superimposed on the magnetic anomaly data in fig. 2 is a shaded relief modelled by using a light-source illumination inclination of 45° and a declination of 45° (i.e. a light-source from the northwest). release of data aeromag 2013 project was marked obtained for free from the ministry of industry and mineral resources in greenland by submitting a form available at the greenland mineral resources portal (http://www. greenmin.gl/). magnetic anomaly maps and geological implications nagssugtoqidian orogen that mainly consists of reworked archaean gneisses with minor supracrustal rocks and several palaeoproterozoic intrusives (bridgwater et al. 1990; kolb in press). palaeogene intrusions and coast-parallel dykes are found in the northern part of the survey area (tegner et al. 1998). aeromag 2013 survey area range in amplitudes between –1318 nt and +3270 nt, with both the most negative and positive values relating to mapped intrusions (figs 2, 3). in the southern part of the survey area, several nne–ssw-trending subparallel linear and positive anomalies (marked i in fig. 2) are anomalies can be traced in the survey area over more than fig. 2. aeromagnetic anomaly map of the survey area in south-eastern greenland covered during the aeromag 2013 project. i–v: magnetic anomalies discussed in the text. i: possible large dykes nne–ssw. ii: magnetic e–w low. iii: ammassalik igneous complex. iv: 1.9–2.2 ga diorite intrusion. v: kruse fjord gabbro complex. http://www.greenmin.gl/ http://www.greenmin.gl/ 65 eral of these anomalies appear to be truncated by a magnetic low (ii; fig. 2) just north of the ammassalik igneous comhave amplitudes in the order of 300 nt, with a magnetisation ing a magnetisation dominated by induced magnetisation or a remanent magnetisation with a direction in the same gen(bridgwater et al. 1990). in order to estimate the depth of the source of the magnetic anomalies, euler deconvolution (reid et al. 1990) was carried out using the standard euler od is based on euler’s homogeneity equation that relates the euler’s equation simultaneously for each grid position within a window and then determines the anomaly position, depth, parameter in the euler deconvolution is the structural index, which enters as an exponential factor corresponding to the a given geometry. for our analysis we used a structural index value of 1, which is suitable for a dyke, and a window size of 2 generally below 400 m. given the uncertainty of the depth area (becker et al. 2009) where the magnetic anomalies are the coast. finally, we note that the kangâmiut dykes found in west greenland, in a similar location on the north atlantic craton (south of the central part of the nagssugtoqidian orogen in west greenland) are much less prominent magnetically (rasmussen & van gool 2000; korstgård et al. 2006) than the linear magnetic anomalies seen in the aeromag 2013 survey area. we tentatively interpret the linear magnetic anomalies as massive dykes. massalik igneous complex marked ii in fig. 2 coincides with a suggested suture zone of the nagssugtoqidian orogen, where the rae craton to the north-east in an oblique wswdirected subduction collided with the north atlantic craton south of the suture zone at c. 1870–1885 ma (kolb in press). by the apparent abrupt termination of several of the positive fig. 3. geological map of the survey area in south-eastern greenland (modi�ed from escher 1990). red: the aeromag 2013 survey area. i–v: magnetic anomalies discussed in the text. 6666 nne–ssw-trending anomalies marked i on fig. 2 and discussed above. north of the ammassalik igneous complex the c. 1.7 ga post-orogenic granodiorite intrusion marked iii in fig. 2 is clearly defined by positive magnetic anomalies with amplitudes up to almost 2000 nt. similarly, the older 1.9–2.2 ga diorite intrusion farther north marked iv in fig. 2 is associated with a strong (c. 2500 nt) positive magnetic anomaly. the strongest magnetic anomalies in the aeromag 2013 survey area are found in the northern part of the survey area and can be related to the palaeogene intrusions in the area. the kruuse fjord gabbro complex marked v in fig. 2 is related to a negative anomaly, hence having a magnetisation dominated by reversely magnetised remanent magnetisation, which is in excellent accordance with an ar-ar isochron age of 48.0 ± 1.2 ma, and emplacement of the intrusive complex during the reverse c21r chron (cande & kent 1995). conclusions in this paper we present the newly released aeromag 2013 survey that adds new and exciting data to the already extensive database of greenland aeromagnetic data. the paper focuses on magnetic anomalies of regional extent, including sub-parallel linear and positive anomalies trending nne– ssw (marked i on fig. 2) that we suggest stem from hitherto undiscovered very large dykes. the suggested location of the suture zone of the nagssugtoqidian just north of the ammassalik igneous complex (ii; fig. 2) is supported by the aeromagnetic data. finally, we note a general good correspondence between the mapped surface geology of the region and the aeromagnetic data; in particular, the intrusions which are clearly discernible. the magnetic data provide a basis for further analysis and modelling of the 3d geometry of the igneous intrusions. many more local anomalies can be identified in the aeromag 2013 dataset but interpretations require further analyses. acknowledgements funding of the aeromag 2013 project was provided by the ministry of industry and mineral resources, government of greenland. thanks are due to eon geosciences inc. for fulfilling all aspects of their contracts in a professional manner. references becker, j.j. et al. 2009: global bathymetry and elevation data at 30 arc seconds resolution: srtm30_plus. marine geodesy 32, 355–371. bridgwater, d., austrheim, h., hansen, b.t., mengel, f., pedersen, s. & winter, j. 1990: the proterozoic nagssugtoqidian mobile belt of southeast greenland: a link between the eastern canadian and baltic shields. geoscience canada 17, 305–310. cande, s.c. & kent, d.v. 1995: revised calibration of the geomagnetic polarity timescale for the late cretaceous and cenozoic.  journal of geophysical research, solid earth 100, 6093–6095. eon geosciences inc. 2013: final survey report. high resolution aeromagnetic survey, southeast greenland aeromag 2013 block, 31 pp. unpublished report, eon geosciences inc., montreal, quebec, canada (in archives of the geological survey of denmark and greenland, report file 23712). escher, j.c. 1990: geological map of greenland, 1:500 000, sheet 14, skjoldungen. copenhagen: geological survey of greenland. kolb, j. in press: structure of the palaeoproterozoic nagssugtoqidian orogen, south-east greenland: model for the tectonic evolution. precambrian research, http://dx.doi.org/10.1016/j.precamres.2013.12.015 korstgård, j.a., stensgaard, b.m. & rasmussen, t.m. 2006: magnetic anomalies and metamorphic boundaries in the southern nagssugtoqidian orogen, west greenland. geological survey of denmark and greenland bulletin 11, 179–184. rasmussen, t.m., thorning, l., riisager, p. & tukiainen t. 2013: airborne geophysical data from greenland. geology and ore, exploration and mining in greenland. 22, 12 pp. rasmussen, t.m & van gool, j.a.m. 2000: aeromagnetic survey in southern west greenland: project aeromag 1999. geology of greenland survey bulletin 186, 73–77. reid, a.b., allsop, j.m., granser, h., millett, a.j. & somerton, i.w. 1990: magnetic interpretation in three dimensions using euler deconvolution. geophysics 55, 80–90. riisager, p., pedersen, m., jørgensen, m.s., schjøth, f. & thorning, l. 2011: dodex – geoscience documents and data for exploration in greenland. geological survey of denmark and greenland bulletin 23, 77–80. tegner, c., duncan, r.a., bernstein, s., brooks, c.k., bird, d.k. & storey, m. 1998: 40ar–39ar geochronology of tertiary mafic intrusions along the east greenland rifted margin: relation to flood basalts and the iceland hotspot track. earth and planetary science letters 156, 75–88. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: pri@geus.dk. geological survey of denmark and greenland bulletin 31, 2014, 1-8 1 geological survey of denmark and greenland bulletin 31 • 2014 review of survey activities 2013 edited by ole bennike, adam a. garde and w. stuart watt geological survey of denmark and greenland ministry of climate, energy and building 22 geological survey of denmark and greenland bulletin 31 keywords geological survey of denmark and greenland, survey organisations, current research, denmark, greenland. cover photographs from left to right 1. a hercules aircraft from the us air force on the greenland ice sheet. photograph: dirk van as. 2. perhaps a future geologist searching for fossils in the faxe quarry. photograph: ole bennike. 3. in 2013 geus celebrated its 125th anniversary. photograph: jakob lautrup. 4. learning about geothermal energy by testing the temperature of the water. photograph: jakob lautrup frontispiece: facing page in 2012 and 2013 geus carried out field work in north greenland (see paper by bojesen-koefoed et al. in this issue). the overturned sedimentary rocks exposed along the margin of the glacier belong to a cretaceous succession in the harder fjord fault zone, wandel sea basin, northern peary land. photograph: jørgen a. bojesen-koefoed. chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, department of geoscience, aarhus university; minik rosing, geological museum, university of copenhagen; finn surlyk, department of geosciences and natural resource management, university of copenhagen scientific editors: ole bennike, adam a. garde and w. stuart watt editorial secretary: jane holst referees (numbers refer to first page of reviewed article): meri-liisa airo, fi (63; morten ahlborn, no (23); jens asger andersen, dk (47); kresten anderskouv, dk (23); anonymous (83, 91); niels balling, dk (75), thue sylvester bording, dk (55); synnøve elvevold, no (95); ida fabricius, dk (27, 55); ola fredin, no (39); høgni kalsø hansen, dk (95); jens havskov, no (75); claus heilmann-clausen, dk (35); björn heincke, de (63); mads huuse, gb (43); christine hvidberg, dk (87); simon toft ingvertsen, dk (47); chris king, gb (31); john korstgård, dk (15, 67); cees laban, nl (43); michael larsen, dk (59); poul-henrik larsen, dk (59); sebastian h. mernild, cl (87); alex mitlehner, gb (35); steffen b. olsen, dk (51); veli-pekka salonen, fi (39); niels schrøder, dk (15); nigel smith, gb (19); inga sørensen, dk (51); leigh stearns, gb (79); lars stemmerik, dk (71); henrik stendal, gl (67); svend stouge, dk (19); vural sander suicmez, dk (27); david sutherland, usa (79); nicolas thibault, dk (31); thomas ulrich, dk (71); frans stephan van buchem, dk (23); jacob clement yde, no (83, 91). illustrations: benny m. schark, jette halskov, willy l. weng, frants v. platen-hallermund and christian brogaard pedersen layout and graphic production: kristian a. rasmussen printer: rosendahls-schultz grafisk a/s, albertslund, denmark manuscripts received: 17 january 2014 – 28 march 2014 final versions approved: 26 february 2014 – 9 may 2014 printed: 25 june 2014 issn 1603-9769 (review of survey activities) issn 1604-8156 (geological survey of denmark and greenland bulletin) isbn 978-87-7871-383-4 citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 31, 98 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk and at www.geus.dk/publications/bull © de nationale geologiske undersøgelser for danmark og grønland (geus), 2014 for the full text of the geus copyright clause, please refer to www.geus.dk/publications/bull mailto:geus@geus.dk 3 44 tanzania ethiopia mozambique zambia malawi indonesia singapore vietnam thailand india tajikistan korea greenland brazil denmark nigeria canada 7 review of survey activities 2013 f.g. christiansen 9 125 years of geological research for society j. fredericia and p. gravesen 15 borehole logs from the precambrian basement on bornholm, eastern denmark: geology and groundwater flow p. gravesen, b. nilsson, p. rasmussen and s.a.s. pedersen 19 the lower palaeozoic shale gas play in denmark n.h. schovsbo, a.t. nielsen and d.l. gautier 23 seismic stratigraphy and sedimentary architecture of the chalk group in south-west denmark c. larsen, j. ineson and l.o. boldreel 27 a novel technique for obtaining representative water samples during co2 core-flooding experiments on chalk at reservoir conditions c. kjøller and j. zuta 31 calcareous nannofossil and foraminifer biostratigraphy of the campanian–maastrichtian chalk of the femern bælt (denmark–germany) e. sheldon, c. morigi and s.d. møller 35 palynological and microfossil biostratigraphy and palaeoecology over the paleocene–eocene transition, femern bælt, northern germany p.a. richardt and e. sheldon 39 ribbed moraines formed during the retreat of the scandinavian ice sheet from eastern himmerland, ne jylland, denmark h. lerche, p.tr. jakobsen and s.a.s. pedersen 43 arctic plant remains of weichselian age from the danish north sea o. bennike, j.o. leth, j.b. jensen, n. nørgaard-pedersen and s. lomholt 47 stormwater management: methods for measuring near-surface infiltration capacity in clayey till b. bockhorn, m.b. jensen and k.e.s. klint 51 a multidisciplinary study of a geothermal reservoir below thisted, denmark m.l. hjuler, h. vosgerau, c.m. nielsen, p. frykman, l. kristensen, a. mathiesen, t. bidstrup and l.h. nielsen 5 tanzania ethiopia mozambique zambia malawi indonesia singapore vietnam thailand india tajikistan korea greenland brazil denmark nigeria canada geus working areas 2013 shown in grey. orange areas are covered in this volume. 55 estimating thermal conductivity from lithological descriptions – a new web-based tool for planning of ground-source heating and cooling c. ditlefsen, i. sørensen, m. slott and m. hansen 59 six years of petroleum geological activities in northeast greenland (2008–2013): projects and a view of the future j.a. bojesen-koefoed, p. alsen and f.g. christiansen 63 aeromagnetic survey in south-eastern greenland: project aeromag 2013 p. riisager and t.m. rasmussen 67 combining exploration and multivariate techniques to detect the bjørnesund west gold occurrence, southern west greenland d.m. schlatter and b. møller stensgaard 71 integrating 3d photogeology with aeromagnetic data as a tool for base-metal exploration in east greenland a. brethes, p. guarnieri and t.m. rasmussen 75 earthquake swarms in greenland t.b. larsen, p.h. voss, t. dahl-jensen and h.p. rasmussen 79 outlet glacier dynamics and bathymetry at upernavik isstrøm and upernavik isfjord, northwest greenland c.s. andresen, k.k. kjeldsen, b. harden, n. nørgaardpedersen and k.h. kjær 83 katabatic winds and piteraq storms: observations from the greenland ice sheet d. van as, r.s. fausto, k. steffen and the promice project team 87 mass loss from an ice-sheet drainage basin in west greenland m.l. andersen, s.b. andersen, l. stenseng, h. skourup, w. colgan, s.s. kristensen, j.p.m. boncori, a.p. ahlstrøm, x. fettweiss, r. forsberg, m. citterio, j.e. box, d. van as and r.s. fausto 91 surface albedo as a proxy for the mass balance of greenland’s terrestrial ice w. colgan, j.e. box, r.s. fausto, d. van as, v.r. barletta and r. forsberg 95 to what extent is denmark vulnerable to mineral supply shortage? p. kalvig, r.j. clausen, n. fold and k. hanghøj 66 7© 2014 geus. geological survey of denmark and greenland bulletin 31, 7–8. open access: www.geus.dk/publications/bull review of survey activities 2013 flemming g. christiansen deputy director 2013 was a very special year for the geological survey of denmark and greenland (geus), which acquired its present name in 1995 when the geological survey of denmark, established in 1888, and the younger geological survey of greenland were merged. on 4 april 2013 geus celebrated its 125th anniversary with a series of presentations by prominent guests and geus scientists followed by a reception and a gala evening. geus’ 125-year long history has provided an overwhelming body of results and experience, and today geus is broader, stronger and more international and collaborative than ever, covering all aspects of which a geological survey can be of use for society. the anniversary is also reflected in this issue of review of survey activities that contains a total of 22 four-page papers, 11 on denmark, nine on greenland and three on broader themes: co2 capture and storage, the recently established center for minerals and materials (mima) housed within the survey – and, of course, a paper dedicated to the 125th anniversary. this first paper highlights the anniversary and introduces the newly published book on the history of geus ‘we – the people down to earth’ (in danish: vi de jordbundne); a book that gives glimpses and highlights of the story of geus spiced with anecdotes on colourful characters, culture and politics. activities in denmark the activities and research in denmark by geus cover many topics within our main programme areas: data, water, energy, mineral resources as well as nature and climate. the island of bornholm has a different geological history from the rest of denmark; a dominance of outcropping basement rocks and palaeozoic sedimentary rocks gives special challenges and possibilities. one paper describes the use of geophysical wire-line borehole logs from the precambrian basement and the mapping of fracture patterns affecting groundwater flow. previous coring results from bornholm have provided important geological input for assessing the unconventional gas resources in the lower palaeozoic shales of denmark. based on several years of collaboration with geus, the u.s. geological survey presented their shale gas estimate in november 2013 suggesting that significant technically recoverable resources may be present in several prospective areas. chalk is a very important rock for denmark as it hosts more than 90% of the danish petroleum reserves in the north sea but it also has many other important uses. one paper is from a pilot project in south-west denmark where seismic stratigraphy and sedimentary architecture of the chalk group has been used in a geothermal evaluation. another paper describes a new technique for important co2 flooding experiments of chalk at reservoir conditions; if these experiments can be up-scaled, their results are potentially crucial to enhanced oil recovery. large infrastructure projects such as tunnels and bridges require detailed geotechnical knowledge of the subsurface, and based on material from new core holes, a more detailed understanding of understudied stratigraphic intervals may be reached. based on cores from the femern bælt, one paper describes a study of the calcareous microfossil and foraminiferal biostratigraphy of the campanian−maastricthian chalk, and another paper focuses on the palynological and microfossil biostratigraphy and palaeoecology of the paleocene− eocene transition. geus is involved in many studies of quaternary and recent geological processes. one paper describes ribbed moraines from eastern himmerland in jylland formed during the retreat of the scandinavian ice sheet. another paper focuses on plant remains from cores in the north sea, and demonstrates that parts of this area constituted an open, treeless, tundra-like environment in the weichselian. a third paper is on stormwater management, where a test site in høje taastrup has been used to evaluate the infiltration capacity of clayey till. based on field experiments, the paper discusses how different methods can be used to measure the infiltration capacity of the sediments. geothermal energy and ground-source heating may be of great importance in many areas in denmark in order to live up to the ambition that by 2050 energy and transport shall be 100% based on renewable energy. over many years, geus has been involved in an increasing number of projects covering different aspects of geothermal energy and groundsource heating. one paper describes a multidisciplinary approach to geothermal reservoir characterisation in thisted, 88 where denmark’s first plant was established already in 1984. the plant now needs to increase its capacity with a new borehole. another paper goes into detail about one of the critical key parameters – thermal conductivity of soils. activities in greenland once again there was a high level of field activities in greenland in 2013 with large field programmes in north greenland and in the gardar province in south greenland. many other field studies were also carried out. results from these large and small projects that are very important for evaluating and marketing the resource potential in greenland and for monitoring climate changes will be presented in the coming years. in this issue, results are presented from other completed and on-going projects. a first paper gives a review of six years of petroleum geological activities in north-east greenland with focus on completed projects, and takes a look at the future. the activities were carried out prior to the preparation of recently completed offshore licensing rounds, and the studies are very important for evaluating key geological risks such as occurrence and quality of petroleum source rocks, understanding critical reservoir intervals and the uplift history. they are therefore crucial for planning of future exploration activities by both the authorities and industry. several papers focus on mapping and evaluation of the mineral potential in greenland. one paper gives details of an aeromagnetic survey that was carried out in southern east greenland in 2013, a region where geus has been active over several years with large completed and planned field programmes. the paper outlines how data can be applied to map regional geological and tectonic features. the project aeromag 2013 has recently been released by the greenland authorities. another paper describes how exploration and multivariate techniques were used to detect a gold occurrence in southern west greenland. a third paper demonstrates how integration of 3d photogeology, aeromagnetic and other geophysical data can be used as a promising tool for base-metal exploration in the remote region of east greenland. earthquake swarms in greenland are a very interesting subject that can now be analysed in greater detail than previously thanks to an increased number of seismic stations. data from older detected swarms have been revisited and two new earthquake swarms have been identified on disko and off south-east greenland. shrinking of the ice sheet and local glaciers in greenland provide a significant contribution to global sea-level rise. understanding the dynamics and calculation of mass loss requires a detailed understanding of climate, glacier configuration and fjord bathymetry. one paper describes a case study from the upernavik isstrøm and upernavik isfjord in north-west greenland. the important monitoring programme of the greenland ice sheet (promice) that was initiated in 2007 continuously supplies crucial data that are used in a number of subsequent projects and in key publications. three papers in this issue use such data from promice. one is on katabatic winds and explanation of piteraq storms and is based on observations from the greenland ice sheet. a second is on mass loss on a basinal scale with its focus on a large area in west greenland. a third paper discusses the possibility of using surface albedo as a proxy for greenland ice mass balance. this is particularly important at the moment to help fill data gaps in gravimetric data from satellites. broader international activities internationally geus works in many different countries with a variety of projects and is also involved in broader thematic studies. the last paper addresses scarcity of mineral raw materials, a theme that has recently been placed high on the political agenda in the us, eu, including in denmark, as well as in asia. as a consequence of this, geus has established the center for minerals and materials (mima) in order to identify and study important raw material chains from source to use. one of the first tasks of mima is to enhance our knowledge of the risk of resource scarcity and the ensuing vulnerability of danish society. geological survey of denmark and greenland bulletin 41, 2018, 87-90 87 rare-earth elements (ree) are considered critical raw materials (crm; ec 2018; us department of the interior 2018) and essential in the technological transformation of the energy sector into carbon-free technologies such as wind turbines, electrified transport and led-lights. the new technologies have led to swiftly expanding markets for ree products, in which china has achieved a monopolistic role in all segments of the ree value chains. political strategies aimed to establish ree supplies outside china are currently being implemented within the eu and in other western countries in order to ensure an adequate future ree supply. however, new ree value chains outside china have not yet materialised. the aim of this paper is to assess whether the global ree supply from present and potential mines can keep pace with the ree demand for the expanding offshore wind energy sector (fig. 1). a successful development of this sector outside china relies on an adequate supply of particularly neodymium (nd) and to some extent praseodymium (pr), terbium (tb) and dysprosium (dy), used in permanent magnets for windmill generators. in 2015, about 82% of the global nd-oxide production was used in the permanent magnets examining the rare-earth elements (ree) supply– demand balance for future global wind power scenarios per kalvig and erika machacek 37 38 mine deposit type pea started/ completed mine development fs started/ completed abandoned/ on hold resource estimate 36 39 8 12 11 9 7 10 22 21 2943 42 40+41 35 34 4527 28 49 26 46 33 32 30 31 25 13 48 17 16 1514 20 18+19 24 44 4723 34 51 2 6 carbonatite iron-oxide-apatite hydrothermal alkaline igneous placer bauxite ion-adsorption clays 1 browns range 2 charley creek 3 dubbo zirconia 4 mount weld 5 nolans 6 yangibana 7 ashram main 8 eco ridge 9 foxtrot 10 kipawa/zeus 11 montviel 12 strange lake 13 bayan obo 14 china minmetals 15 china minmetals 16 china minmetals 17 chinalco 18 ganzhow mining group 19 guangdong rising nf 20 xiamen tungsten 21 kringlerne 22 kvanefjeld 23 chavara 24 manavalakurichi 25 kutessay ii 26 ngualla 27 kangankunde 28 songwe hill 29 lofdal 30 khibiny 31 lovozero 32 norra kärr 33 aksu diamas 34 ngualla, tanz. 35 wigu hill 36 bear lodge 37 bokan 38 mountain pass 39 round top 40 glenover (rsa) 41 steenkampskraal 42 zandkopsdrift 43 buena norte 44 lahat 45 tantalus 46 gakara 47 dong pao 48 tomtorskoye 49 nkombwa hill geus/mima april 2018 fig. 1. global rare-earth element mines and advanced exploration projects. fs: feasibility study. pea: pre-economic assessment. © 2018 geus. geological survey of denmark and greenland bulletin 41, 87–90. open access: www.geus.dk/bulletin http://www.geus.dk/bulletin 8888 production (adamas 2016). here we evaluate the future supply and demand situations for nd, pr, tb and dy in the global wind energy sector in the form of three scenarios, one for 2020 and two for 2030 based on high and low demand. the balance is discussed. our assessment reflects the challenge caused by limited insight into the ree supply chains inside china, and the figures presented in this paper are therefore only indicative. scenarios for future global ree demand of the wind energy sector in 2016, the global wind energy council reported a total global wind power capacity of 487 gw (gwec 2016), of which the offshore capacity amounted to 3% according to the global status report 2017 for renewables. due to the otherwise very high maintenance costs of the offshore wind energy, it depends in part on direct-drive and hybrid wind turbine technologies that use ree-based permanent magnet (pm) or high-temperature superconducting (hts) generators (barteková 2016). this is why the offshore wind energy sector is the focus of this study. our scenarios solely deal with technologies using permanent magnets in which nd is vital. barteková (2016) specifies the consumption of total rare-earth element oxides (treo) for the individual types of magnets used in wind turbines and indicates the individual ree used. the reo content per magnet varies by generator design between c. 23–35%, highest in the permanent magnet of a direct-drive synchronous generator (pmsg-dd), and lowest in hybrid single and multistage synchronous generators (pmsg-sg and pmsgmg). the relative magnet weight proportions of the four reos are about 95% nd, 4% pr, 0.99% tb and 0.01% dy (barteková 2016). in order to estimate the future reo-demand of the sector, the following assumptions about the wind energy technology are made based on barteková (2016, p. 158), who reported that in 2014 the ree-based permanent magnet technology accounted for 4% of the offshore wind technology, equally divided between direct-drive and hybrid generator designs. for our 2016 baseline scenario, we (i) increased this share to 5%, (ii) maintained the equal split between direct-drive and hybrid generator technologies, and (iii) subdivided the hybrid generator designs equally into single and multistage gearbox designs. further, we assume that in 2016, china held 8% and the rest of the world (row) 92% of the global offshore wind capacity. from the 2016 baseline scenario, we developed three scenarios: (1) the reo use in the total global offshore wind energy capacity in 2020, and (2, 3) low and high forecasts of the same in 2030. we set the regional offshore shares to 85% for row and 15% for china. see details in table 1. pr nd tb dy pr nd tb dy pr nd tb dy pr nd tb dy 2016 2020 2030-low 2030-high     row 104 2.5 26 0.26 267 6.3 66 1 589 13.9 146 1 986 23.4 244 2 9 ton ton tonkton ton ton tonkton ton ton tonkton ton ton tonkton 215 2 0.02 23 551 6 0 104 2.5 26 0 174 4.1 43 0china 0 5000 0 10 000 15 000 20 000 25 000 ton a     china 0 200 400 600 800 1000 1200 ton pr 4 0 90 8 1 0 0 0 10 1 230 20 2 0 0 0 0 0 45 8 0 0 11 2 1060 187 27 5 631 111 7 1 row nd tb dy pr nd tb dy pr nd tb dy pr nd tb dy 2016 2020 2030-low 2030-high b fig. 2. rare-earth metal demand by wind energy deployment in china and the row in 2016, and forecasts for 2020 and 2030. a: assuming that all offshore wind technology is centred on ree-based permanent magnets. b: assuming that only a small share of offshore wind technology uses reebased permanent magnets, and differentiating between varying ree uses per pm generator technology. note: both figures show cumulated (forecasted) individual ree use by examined wind energy technology in the respective year. row: the rest of the world. table 1. forecast scenarios for installed global offshore wind energy capacity in 2020 and 2030 forecast 2020 739 gw 5% = 37 gw 5% = 1.85 gw forecast 2030 low 1260 gw 7% = 88 gw 10% = 8.80 gw forecast 2030 high 2110 gw 7% = 148 gw 10% = 14.80 gw ree permanent magnets installed capacity offshore capacity %assumptions, this survey from gwec (2016a) global totalwind energy capacity scenarios global offshore 89 our scenarios are based on the estimates by gwec (2016) for 2020, 2030-low, and 2030-high global wind energy capacity. it is important to note the large variability in the underlying assumptions. the most significant parameter is the share of ree-based permanent magnets deployed in offshore installations. for instance, if all offshore technology capacity would employ ree-based permanent magnet technology and 75% of this was installed with the highest ree-using direct-drive design, this would result in the consumption 16 400–27 500 tons nd-oxide in 2030 ( fig. 2a). in contrast, the wind-turbine sector will demand only about 740–1250 tons nd-oxide (fig. 2b) for the alternative technology split-up outlined in table 1. scenarios for the future global ree supply in 2016, the primary global production of total rare-earth oxides (treo) amounted to 129 000 tons, of which china produced 83%, australia 11%, russia 2%, brazil 1%, india 1% and malaysia, thailand, and vietnam still less (usgs 2017). although the treo supply figures for 2016 reported by different sources are rather similar (usgs 2018: 23 680 tons; adamas 2016: 24 377 tons), there are major discrepancies at national level. this partly stems from the assumed contributions of the non-reported market which may account for 25–30% (roskill 2016) and from uncertainties pertaining to production and smelting quota (adamas, personal communications 2018). this study applies the figures for for 2016 from usgs (2018), and contribution from nonreported production is not considered. the 2016 supplies of pr, nd, dy and tb from china and seven row countries are shown in figs 3a, b. the official chinese production quota for 2016 was set to 105 000 tons treo (machacek & kalvig 2017); our estimate of the regional reo production is based on kingsnorth (2016) and shown in fig. 3a. given that no scheduled production quotas for 2020 and 2030 are available, our china 2020 and 2030 supply scenarios are arbitrarily set to an increase of 5% p.a., reflecting the anticipated growth in demand (dutta et al. 2016), the potential for higher capacity on existing plants, as well as continued efforts to transform the informal sector into to a formal one. the row supply in the 2020 and 2030 scenarios is developed as follows: of the recorded 320 ree exploration projects outside china, 99 are reported to be active (s&pdatabase 2017). our search revealed that 31 of these projects, located in 12 countries, have reached an advanced stage (fig. 1). these 31 projects are divided into four classes of development, which are in turn translated into expected production start-ups in 2020, 2025, 2030 and 2035. the estimates of relative reo grades and targeted production of pr, nd, tb and dy are based on tmr (2015) and company data. where relative grade data are not available, our estimate is based on the actual ree mineralogy. according to these data, new reo productions in australia, usa, and vietnam are expected in 2020. in 2030, reo will also 5000 10 000 15 000 20 000 25 000 30 000 35 000 pr a nd tb dy pr nd tb dy pr nd tb dy 2016 2020 (5% growth p.a.) 2030 (5% growth p.a.) yunnan hunan fujian guangdong guangxi jiangxi shandong sichuan inner monglia to n 46 69 17 4 23 28 4 88 6 56 75 21 1 78 34 5 56 2 17 54 10 77 9 24 4 34 4 96 1000 0 2000 3000 4000 brazil vietnam thailand malaysia india russia australia b pr nd tb dy pr nd tb dy pr nd tb dy 2016 2020 2030 to n 10 49 32 74 62 33 1 10 49 37 24 62 33 1 10 49 37 24 62 33 1 5000 0 10 000 15 000 20 000 25 000 vietnam usa turkey tanzania sweden south africa namibia malawi greenland canada burundi australia c 2016 2020 2030 pr nd tb dy pr nd tb dy pr nd tb dy 0 0 0 0 23 2 76 0 32 20 5 72 04 24 9 49 37 7 21 59 to n fig. 3. current and future ree production scenarios. a: current and forecasted chinese production of praseodymium, neodymium, terbium and dysprosium in seven regions. forecast assumes a general 5% annual increase. b: current and forecasted production of praseodymium, neodymium, terbium and dysprosium in row. forecast assumes a static production. c: forecasted production of praseodymium, neodymium, terbium and dysprosium from 31 advanced ree projects outside china. 9090 be supplied by canada, greenland, malawi, namibia, south africa, sweden, tanzania and turkey (figs 1, 3c). the scenarios indicate row-treo productions of 6920 tons in 2020 and 154 075 tons in 2030 (fig. 3c). discussion and summary we demonstrate here that the level of detail applied to estimates of the future use of different types of generators and their relative shares allow for great variance in the current and forecast ree demand by the wind energy sector. against this background, our row scenario for 2020 points to a nd demand by the wind energy sector within a wide range of 230–6300 tons (fig. 2), while our row supply forecast for nd-oxide from both current and new mines is around 4500 tons. a top-down approach indicates that the ree-based permanent magnets for the wind sector absorb 10%, equivalent to about 4000 tons treo (lucas et al. 2015). adamas (2016) estimates that nd-oxide accounts for c. 73% of the treo in the permanent magnets for the wind sector. in effect this means that in 2020, the total nd-oxide supply for the wind energy sector would be roughly 3000 tons nd-oxide, i.e. in the middle range of the forecasted global demand for this purpose. for the 2030 low scenario and with ree technology applied in some but not all offshore technology, the global nd-oxide demand by the wind sector is forecasted to c. 740 tons and for the 2030 high scenario close to 1250 tons. if it is assumed that all offshore technology will draw on ree use, these figures increase to 16 400 and 27 500 tons, respectively. our global supply forecast of nd-oxide in 2030 from current operations is about 35 000 tons from china and 3700 tons from row, to which advanced row ree projects could contribute an additional c. 25 000 tons if all projects go into production. if the 10% share of the ree-based permanent magnet sector as well as a stable demand for ree-based permanent magnets from the wind sector are maintained, about 450 tons nd-oxide could be made available by row suppliers in 2020, and 2900 tons in 2030. however, in the scenarios based on generator technologies that consume a higher percentage of ree-based magnets, the estimated row supply is inadequate and an additional nd-oxide supply will be required, e.g. from chinese ree operators. this study shows that there are currently significant uncertainties in trying to determine both the current ree demand and supply in the wind sector and in building scenarios for 2020 and 2030. given that nd-dy permanent magnets represent a fast-growing sector, there is a need to establish a comprehensive, research-based and harmonised framework for precise estimates of the future supply and demand scenarios for ree-based permanent magnets. references adamas 2016: rare earth market outlook: supply, demand, and pricing from 2016 through 2025 report, 822 pp. adamas intelligence. barteková, e. 2016: the role of rare earth supply risk in low-carbon technology innovation. in: borges de lima, i. & leal filho, w. (eds): rare earths industry: technological, economic, and environmental implications. amsterdam: elsevier, 437 pp. dutta, t., kim, k-h, uchimiya, m., kwon, e.e., jeon, b-h, deep, a. & yun, s-t. 2016: global demand for rare earth resources and strategies for green mining. environmental research 150, 182–190. ec 2018: report on critical raw materials and the circular economy, 69 pp. commission staff working document. swd(2018) 36 final. gwec 2016: global wind energy outlook 2016. fried, l. et al. (eds), 44 pp. gwec and institute for sustainable futures, university of technology sydney. kingsnorth, d.j. 2016: curtin university publication on the rare earth industry in 2016. lucas, j., lucas, p., le mercier, t., rollat, a. & davenport, w. 2015: rareearth-based permanent magnets preparation and uses, 231–249. in: rare earths. amsterdam: elsevier. machacek, e. & kalvig, p. (eds) 2017: road map for ree material supply autonomy in europe. eurare, european ree market survey (component d1.2). 141 pp + appendix. roskill informations services ltd. 2016: rare earths: global industry, markets and outlook to 2026, 396 pp. roskill reports on metals and minerals, 16th revised edition. s&p-database 2017: s&p global market intelligence metals & mining database 2017. technology metals research (tmr) 2015: advanced rare-earth projects: http://www.techmetalsresearch.com/metrics-indices/tmr-advancedrare-earth-projects-index/ (accessed 9 february, 2018). us department of the interior 2018: draft list of critical minerals. feb 16, notices. federal register 83(33), 7065–7068. usgs 2018: mineral commodity summaries. us geological survey. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: pka@geus.dk. http://www.techmetalsresearch.com/metrics-indices/tmr-advanced-rare-earth-projects-index/ http://www.techmetalsresearch.com/metrics-indices/tmr-advanced-rare-earth-projects-index/ mailto:pka@geus.dk e2019430106-01 the brackish water baltic sea and the more saline kattegat in the north are connected by three straits, lillebælt, storebælt and øresund (fig. 1). storebælt (the great belt) is the deepest and widest of the straits. the strait is characterised by deeply incised channels that are partly filled by sediments. the water depth in major parts of storebælt is about 20 m, though in some areas the channels are more than 50 m deep. the formation of the channels has been subject to discussion. andersen (1927) suggested that the channels formed due to strong currents that are still active today or by fluvial erosion during the so-called continental period (fastlandstiden) in the early holocene. at this time, the relative sea level in the region was lower than at present and a huge lake, the ancylus lake, which occupied the baltic basin, may have drained via storebælt. andersen dismissed the idea that the channels were formed by subglacial erosion by meltwater during the last deglaciation. more recently, mathiassen (1997) interpreted some of the deposits in the channels as late glacial, a viewpoint followed by bennike et al. (2004). however, the age of the late glacial deposits in the channels are poorly constrained. the first studies of sediment cores from storebælt were carried out by krog (1973), winn (1974) and mathiassen (1997), but these studies concentrated on the holocene development from mires to lakes to brackish and marine environments. wiberg-larsen et al. (2001) documented the presence of early holocene river deposits. here we report on some new ages of macrofossils from late glacial deposits in the storebælt channels. in the late 1970s, the danish state began to map deposits of sand and gravel in storebælt and several potential aggregate resources were identified east of romsø. in this part of storebælt there are two parallel south–north-orientated incised channels. in 2017, new vibrocores were collected by the geological survey of denmark and greenland (geus) as part of a mapping programme of submarine aggregates the channels in storebælt, denmark: implications of new radiocarbon ages ole bennike*1, niels nørgaard-pedersen1 and jørn bo jensen1 geus bulletin is free to individuals and institutions in electronic form. the author(s) retain copyright over the article contents. research article | open access geus bulletin vol 43 | e2019430106 | published online: 24 june 2019 https://doi.org/10.34194/geusb-201943-01-06 10°e 14°e 56°n 57°n sweden denmark germany kattegat storebælt femer bælt baltic sea kielerkieler lillebælt øresundfig. 2 50 km jylland fyn 55°n fig. 1. map of denmark and the surrounding area showing the location of the study area. fig. 2. bathymetry of the storebælt region showing the locations of the cores discussed in this paper. water depth: 0 to c. 50 m. water depth at core site 544001-1 is 38.5 m, and at core site 544002-1 it is 33.6 m. 10°40'e 11°10'e 55°40' 55°30'n 5 km 544001-1 544002-1 544001-1 fyn sjælland 544002-1 romsø https://doi.org/10.34194/geusb-201943-01-06 e2019430106-02 for the danish environmental protection agency. coring positions were selected from interpretations of shallow seismic data acquired during the initial mapping projects. two of the new cores from storebælt contained remains of plants that are typical of late glacial deposits from denmark. three samples were submitted for radiocarbon dating, and here we report on the results. methods an innomar parametric sub-bottom profiler was used to acquire high-resolution seismic data. vibrocoring was carried out with a 6 m long, 10 cm diameter vibrocorer (vkg-6) from the vessel skoven. differential gps was used to determine coring locations. the cores were cut into 1 m long sections onboard the ship and sent to geus. in the laboratory, the core was split, photographed and described. half of the cores were archived and the other half were subsampled. samples for palaeoecological analyses and radiocarbon dating were wet sieved and the residue left on the sieves was analysed using a dissecting microscope. ages were determined by accelerator mass spectrometry (ams) radiocarbon dating and calibrated to calendar years before present (1950 ce) according to the intcal13 data (table 1). results core 544001-1 was collected on the western flank of the western storebælt channel (fig. 2). a shallow seismic east– west profile obtained at the coring site shows thin holocene marine deposits underlain by late glacial deposits (lomholt et al. 2017). the core is 557 cm long and contains fine-grained sand overlain by silt with granules. these lower units do not contain any macrofossils (fig. 3). then follows an almost 400 cm thick unit of heterolithic laminated mud and clay with plant and animals remains. plant remains include the land plants betula nana, betula pubescens, dryas octopetala, selaginella selaginoides and distichium sp., the reed plant scirpus lacustris and the water plants ranunculus hyperboreus, hippuris vulgaris, eleocharis palustris, potamogeton natans, p. perfoliatus, p. praelongus, stuckenia pectinata and s. filiformis. animal remains comprise piscicola geometra, candona sp., cytherissa lacustris, valvata cristata, v. piscinalis, pisidium sp. and cristatella mucedo. the assemblage is typical of late glacial lake deposits in the region, with the exception of scirpus lacustris and stuckenia pectinata (jensen 1985; bennike et al. 2004). the warmth-demanding scirpus lacustris (schoenoplectus lacustris) is rare in late glacial deposits, though it has been recorded in sediments from the arkona basin in the south-western baltic sea, where it was dated to c. 12 800 cal. years bp, corresponding to the end of the allerød chronozone. to our knowledge, there are no previous finds of stuckenia pectinata from late glacial deposits in denmark. these plants probably spread northwards by rivers in storebælt during the bølling and allerød chronozones. a sample of betula nana twigs from a depth of 447 cm was dated to 14 070–14 560 cal. years bp, corresponding to the bølling chronozone, and a sample of scirpus lacustris was dated to 13 094–13 315 cal. years bp, corresponding to the allerød chronozone (table 1). the upper 1.5 m consists of marine bioturbated clay and mud with shells of marine gastropods, bivalves and barnacles, such as turritella communis, arctica islandica and balanus crenatus. core 544002-1 is 500 cm long and contains sandy diamicton at its base, which is interpreted as glacial till (figs 3, 4). the till is overlain by laminated mud and homogenous finegrained sand with plant and invertebrate remains. remains of land plants are dominated by betula nana but also include dryas octopetala, salix sp., empetrum nigrum, rumex acetosella and selaginella selaginoides. water plants include ranunculus sect. batrachium sp., menyanthes trifoliata, stu ckenia pectinata and s. filiformis. the fauna includes candona sp., cytherissa lacustris, lepidurus sp., gyraulus rossmaessleri, pisidium sp. and cristatella mucedo. a sample of betula nana and dryas octopetala remains was dated to 13 130–13 340 cal. years bp, which corresponds to the allerød chronozone. the flora and fauna are typical of late glacial deposits in the region, except for gyraulus rossmaessleri, which table 1. new radiocarbon ages from storebælt, denmark * below core top. † calibrated to calendar years before present according to the intcal13 dataset. 11 374 ± 54 9294 ± 33 core n. lat. e. long. laboratory species depth age (14c calibrated no. no. b.c.t. (cm)* years bp) age (years bp)† 12 310 ± 40 11 405 ± 44 9455 ± 38 544001-1 55.571° 10.828° aar-29105 scirpus lacustris 201 13 094–13 315 544001-1 beta-481723 betula nana 447 14 070–14 560 544002-1 55.538° 10.823° aar-29106 phragmites australis 214 10 433–10 561 544002-1 aar-29107 phragmites australis 229 10 604–10 741 544002-1 beta-481724 b. nana, dryas octopetala 375 13 130–13 340 e2019430106-03 is rarely recorded, and stuckenia pectinata as mentioned above. the heterolithic sediments are overlain by a layer of homogenous fine-grained sand with betula nana remains. the minerogenic sediments are overlain by a thin layer of peat and gyttja with a rich flora that includes pinus sylvestris, betula sect. albae sp., populus tremula, cladium mariscus, stachys palustris, ceratophyllum demersum, najas marina and a fauna with bithynia tentaculata and perca fluviatilis. we also found a fragment of a lower jaw of sorex minutus (pygmy shrew). the fossil assemblage is typical of early holocene peat and gyttja deposits in the region (bennike et al. 2004), but sorex minutus is new to the early holocene fauna of denmark (aaris-sørensen 2009). the lower part of the peat was dated to 10 604–10 741 cal. years bp and the upper part to 10 433–10 561 cal. years bp (table 1). the upper 2 m of the sediment core consists of bioturbated mud with shells of marine species such as ammonia beccarii, balanus crenatus, tritia reticulata, nucula tenuis, arctica islandica, abra alba, corbula gibba and echinocardium cordata. fig. 3. sedimentological logs based on sediment cores from storebælt. radiocarbon ages are calibrated to calendar years bp (mean probability ages). vf: very fine-grained. f: fine-grained. m: medium-grained. c: coarse. vc: very coarse. 1 0 2 3 4 5 cla y sil t vf sand f m c vcd ep th (m ) 0 d ep th (m ) li th ol og y li th ol og y 544001-1; water depth 38.5 m 14 244 cal. years bp 1 2 3 4 544002-1; water depth 33.6 m 13 242 cal. years bp 13 215 cal. years bp 10 499 cal. years bp 10 691 cal. yrs bp glacial till homogenous marine holocene mud with shells homogenous fine-grained sand peat homogenous marine holocene clay with shells alternating layers of mud and clay homogenous fine-grained sand silt with pebbles layered mud homogenous marine holocene mud with shells cla y sil t vf sand f m c vc fig. 4. selected part of sub-bottom innomar profile, with the position of sediment core 544002-1 indicated. the red line separates till from late glacial non-marine sediments. the blue line separates late glacial sediments from marine sediments. d ep th b el ow se a lev el (m ) 40 42 44 38 36 34 32 30 core 544002-1 late glacial marine till 500 mnnw sse e2019430106-04 discussion the old age of c. 14 200 cal. years bp from core 544001-1 is noteworthy, because only a few plant remains of this age have previously been radiocarbon dated from late glacial deposits in denmark. active glacier ice disappeared from the region between 18 000 and 17 000 years bp (houmarknielsen et al. 2012). stagnant ice lingered on in many parts of denmark and at the same time, temperatures were low in the region until the abrupt warming at the start of the bølling chronozone at 14 700 years bp. the widespread existence of stagnant glacier ice and low temperatures likely delayed the migration and spread of terrestrial plants and animals in the region. in core 544002-1, a layer of sand is found below the peat deposit. such sand layers are commonly found below early holocene peat layers in the region. in some cores the sand layers contain in situ roots that can be related to the peat layers. in a few cases, the sand layers contain an increasing amount of plant remains in the upper part, and the boundary between the sand layer and the peat layer is often gradual. however, in most cases the sand layers contain no plant macrofossils and the sand–peat boundary is sharp. the sand layer probably reflects a low stand, and the following peat formation can be seen as a consequence of increasing temperatures and more stable soil conditions at the beginning of the holocene. we interpret the sand layer as late glacial because it contains some remains of betula nana, but usually such sand layers do not contain plant macrofossils (bennike et al. 2004). conclusions the radiocarbon ages discussed in this paper indicate that the deeply incised channels in storebælt were formed by meltwater during the last deglaciation, as the ice margin receded southwards and huge amounts of meltwater flowed northwards. the channels were probably eroded by water below the ice, near the ice margin. the fossil flora and fauna are typical of late glacial deposits in the region, except for the warmth-demanding reed plant scirpus lacustris and the water plant stuckenia pectinata. these species likely colonised the storebælt area due to northward flowing rivers and grew in the area during the warm bølling and allerød interstadials. the ages show that the late glacial deposits in the incised channels are at least in part of bølling and allerød age. acknowledgements this study was supported by geocenter denmark. kristian gregersen from the zoological museum in copenhagen kindly identified the jaw fragment of sorex minutus. we are grateful to journal referees jacob yde and jakob qvortrup christensen for comments on the manuscript. references aaris-sørensen, k. 2009: diversity and dynamics of the mammalian fauna in denmark throughout the last glacial–interglacial cycle, 115–0 kyr bp. fossils and strata 57, 59 pp. andersen, s.a. 1927: storebælt i nutid og fortid. geologiska föreningens i stockholm förhandlingar 49, 427–437. https://doi. org/10.1080/11035892709444582 bennike, o., jensen, j.b., lemke, w., kuijpers, a. & lomholt, s. 2004: lateand postglacial history of the great belt, denmark. boreas 33, 18–33. https://doi.org/10.1111/j.1502-3885.2004.tb00993.x houmark-nielsen, m., linge, h., fabel, d., schnabel, c., xue, s., wilcken, k.m. & binnie, s. 2012: cosmogenic surface exposure dating the last deglaciation in denmark: discrepancies with independent age constraints suggest delayed periglacial landform stabilization. quaternary geochronology 13, 1–17. https://doi.org/10.1016/j.quageo.2012.08.006 jensen, h.a. 1985: catalogue of lateand post-glacial macrofossils of spermatophyta from denmark, schleswig, scania, halland, and blekinge dated 13,000 b.p. to 1536 a.d. danmarks geologiske undersøgelse, serie a 6, 95 pp. krog, h. 1973: the early holocene development of the store belt as reflected in a former fresh water basin. danmarks geologiske undersøgelse, årbog 1972, 37–47. lomholt, s., leth, j.o., nørgaard-pedersen, n., witt, n.h. & bennike, o. 2017: rapportering af 150 boringer udført i spekulative ressourcer. danmarks og grønlands geologiske undersøgelse rapport 2017/34, 248 pp. mathiassen, d.r. 1997: the changing landscapes of the storebælt from the retreat of the ice to the sea flood. in: pedersen, l., fischer, a. & aaby, b. (eds): the danish storebælt since the ice age: man, sea and forest. a/s storebæltsforbindelsen, copenhagen, 22–28. wiberg-larsen, p., bennike, o., jensen, j. b. & lemke, w. 2001: trichoptera remains from early holocene river deposits in the great belt, denmark. boreas 30, 299–306. https://doi.org/10.1111/j.1502-3885.2001. tb01049.x winn, k. 1974: present and postglacial sedimentation in the great belt channel (western baltic). meyniana 26, 63–101. how to cite bennike, o., nørgaard-pedersen, n. & jensen, j.b. 2019: the channels in storebælt, denmark: implications of new radiocarbon ages. geological survey of denmark and greenland bulletin 43, e2019430106. https://doi.org/10.34194/geusb-201943-01-06 *corresponding author: ole bennike | e-mail: obe@geus.dk 1 geological survey of denmark and greenland (geus), c.f. møllers allé 8, dk-8000, aarhus c, denmark. https://doi.org/10.1080/11035892709444582 https://doi.org/10.1080/11035892709444582 https://doi.org/10.1111/j.1502-3885.2004.tb00993.x https://doi.org/10.1016/j.quageo.2012.08.006 https://doi.org/10.1016/j.quageo.2012.08.006 https://doi.org/10.1111/j.1502-3885.2001.tb01049.x https://doi.org/10.1111/j.1502-3885.2001.tb01049.x https://doi.org/10.34194/geusb-201943-01-06 mailto:obe@geus.dk e2019430101-01 between 1850 and 2006 global mean sea level rose by 24 ± 18 cm. it is projected to rise a further 52 ± 21 cm under the representative concentration pathway (rcp) 4.5 scenario, which approximates the carbon emissions reductions of the ‘paris agreement’ climate pathway. it is projected to rise 74 ± 28 cm under the rcp8.5 scenario, which represents a ‘business-as-usual’ climate pathway (box & colgan 2017). these rates of recent and future sea-level rise are faster than those reconstructed for previous warm intervals, such as the medieval climatic optimum (c. 1000 to 1400 ce) and the holocene thermal maximum (c. 7000 to 3000 bce) (gehrels & shennan 2015). moreover, palaeo reconstructions indicate a global sea-level sensitivity of two metres per degree of warming (levermann et al. 2013). the forces driving global sea-level change are complex. the global sea-level budget includes the transfer of land ice into the ocean, thermal expansion of seawater, changes in land water storage, and changes in ocean basin volume (church et al. 2013). at the local scale, the evolving planetary gravity due to shifting water and ice masses, shifting oceanic and atmospheric currents and persistent tectonic and glacial isostatic adjustment processes can also be important. sea-level changes around the globe are therefore far from uniform (jevrejeva et al. 2016). here, we highlight the value of combining palaeo reconstructions of sea level, the measured tide gauge record, and projections of future sea level. this allows us to understand local sea-level changes from the recent past in the context of global projections for the near future (0 to 2100 ce). we explore the strong differences in local sea-level histories and future projections at three danish cities: skagen and esbjerg, as they have contrasting glacio-isostatic adjustment histories, and copenhagen, where we also compare local and global drivers of present-day sea-level rise based on previously published research. data we employ the standardised permanent service for mean sea level annual tide gauge records at copenhagen (psmsl site 21), esbjerg (psmsl site 80) and skagen (psmsl site 89) since c. 1880 (holgate et al. 2013). while psmsl data are formatted as sea-level elevation relative to the geoid, we instead express local sea-level elevations as relative to the 1901–1950 baseline elevation throughout this study (fig. 1). for each city, we characterise a centennial (1900–1999) rate of sea-level change using a linear trend to the annual psmsl data. we estimate uncertainty in this centennial sea-level trend using a monte carlo envelope that assumes ± 10 cm uncertainty in annual elevations. the dating of raised beach sequences, wind-blown sand deposits, and salt-marsh sediments has permitted sea level to be reconstructed since the last glaciation at skagen (hauerbach 1992; clemmensen et al. 2001) and esbjerg (gehrels et al. 2006; szkornik et al. 2008). at esbjerg, the palaeo records of relative sea level overlap with the observed record of tide gauge data, revealing some discrepancies between the two datasets during this overlapping period. for example, gehrels et al. (2006) suggest palaeo-sea-level at esbjerg in the 1880s was –24 cm below the c. 2000 level, while the tide gauge record suggests it was –5 cm below the 1901–1950 baseline. some of this apparent discrepancy is likely attributable to differing baseline periods in the two datasets. to minimise such discrepancies, we shift the relative sea levels reported in gehrels et al. (2006) and szkornik et al. (2008) by +19 and –8 cm, respectively. at skagen, there is no overlap of measured tide gauge data and reconstructed sea levels. we therefore assume the relative sea levels reported in hauerbach (1992) and clemmensen et al. (2001) are characteristic of the 1901–1950 reference period. analogous palaeo reconstructions of sea-level elevation are not readily available in the immediate vicinity of copenhagen, reflecting the relative paucity of palaeo data within urbanised sjælland. sea-level rise in denmark: bridging local reconstructions and global projections william colgan*1, jason e. box1, sofia ribeiro1 and kristian k. kjeldsen1 geus bulletin is free to individuals and institutions in electronic form. the author(s) retain copyright over the article contents. research article | open access geus bulletin vol 43 | e2019430101 | published online: 01 july 2019 https://doi.org/10.34194/geusb-201943-01-01 https://doi.org/10.34194/geusb-201943-01-01 e2019430101-02 the projections of future sea-level rise are based on simulations of global sea-level budget terms under the rcp4.5 and rcp8.5 climate pathways (box & colgan 2017). we translate this 1850–2100 global eustatic sea-level budget into local sea-level budgets by applying a linear trend to the global budget that makes it fit the linear trend of a local budget during the 20th century (1900–1999). this yields global-to-local scaling terms of 0.4 mm/year at esbjerg, –1.8 mm/year at skagen and 1.1 mm/year at copenhagen. these terms capture the site-specific processes causing deviations from the global mean – including the net effects of glacioisostatic adjustment and persistent changes in atmospheric and oceanic currents – during the 20th century. these linear scaling terms are also applied to the projections. calendar year 1900 1950 2000 2050 2100 se a le ve l r el at ive 1 90 119 50 (m ) -0.2 0 0.2 0.4 0.6 0.8 1 1.2 calendar year 0 500 1000 1500 2000 se a le ve l r el at ive 1 90 119 50 (m ) -0.5 0 0.5 1 1.5 2 2.5 3 3.5 4 4.5 clemmensen2001 hauerbach1992 holgate2013 box2017:rcp4.5 box2017:rcp8.5 b calendar year 1900 1950 2000 2050 2100 se a le ve l r el at ive 1 90 119 50 (m -0.2 0 0.2 0.4 0.6 0.8 1 1.2 calendar year 0 500 1000 1500 2000 se a le ve l r el at ive 1 90 119 50 (m ) -3 -2.5 -2 -1.5 -1 -0.5 0 0.5 1 1.5 2 szkornik2008 gehrels2006 a holgate2013 box2017:rcp4.5 box2017:rcp8.5 fig. 1. sea level, relative to the 1901–1950 period, at esbjerg (a) and skagen (b) between 0 and 2100 ce, derived from palaeo reconstructions (hauerbach 1992; clemmensen et al. 2001; gehrels et al. 2006; szkornik et al. 2008), tide gauge measurements (holgate et al. 2013) and projections (box & colgan 2017). for palaeo reconstructions, uncertainty is depicted with xand y-whiskers. for projections, uncertainty is depicted with dashed lines bounding solid line best estimates. linear best fit and associated uncertainty is shown for the tide gauge records. e2019430101-03 past millennial-scale sea-level change in the past 2000 years, local relative sea level has risen 1.5 ± 0.5 m (0.8 ± 0.3 mm/year) at esbjerg and fallen 4.0 ± 0.5 m (2.0 ± 0.3 mm/year) at skagen (table 1). these contrasting sea-level histories are primarily due to local differences in glacio-isostatic adjustment. at skagen, the earth’s crust is still rebounding upwards following the relatively rapid removal of the scandinavian ice sheet during the last glaciation c. 17  000 years ago (morén et al. 2018). at esbjerg, the earth’s crust is still sinking due to the collapse of the crustal forebulge that once ringed the scandinavian ice sheet (stuhne & peltier 2015; fig. 2). local relative sea-level rise reflects the net effect of changes in land and ocean elevation. during the 20th century, skagen was rebounding faster than eustatic sea level was rising. as a result, while global average sea level rose 1.5 ± 0.4 mm/year during 1900–1999 (box & colgan 2017), the relative sea level measured at skagen fell 0.3 ± 0.4 mm /year (holgate et al. 2013). at esbjerg, sea level did rise, but less than the global mean (1.1 ± 0.4 mm/ year). these local departures from the global pattern reflect our global-to-linear scaling terms described above. assessing the magnitude and spatial distribution of recent sea-level rise across denmark therefore requires – among other things – constraining present-day glacio-isostatic adjustment rates resulting from deglaciation following the last glacial period. present-day sea-level rise components land ice was responsible for c. 51% of global mean sea-level rise during 2004–2010 (box & colgan 2017). as large ice and water masses shift around the planet, they modify the planetary gravity field. where land ice diminishes under climate change, local gravitational fields weaken and nearby sea level falls. in this process, the ocean water previously held near land ice is redistributed to raise distant sea levels, elsewhere. in this way, greenland land ice contributes four times as much to global mean sea-level rise (c. 24%) than it does to local sea-level rise at copenhagen (c. 6%; fig. 3). conversely, antarctic land ice is slightly more important to local sea-level rise at copenhagen (c. 11%) than the global mean (c. 7%). this is because copenhagen lies within the gravitational weakening anomaly associated with greenland, but lies outside the analogous gravitational weakening anomaly associated with antarctica (larour et al. 2017). notably, while scandinavian land ice contributes just c. 1% global mean sea-level rise, it actually provides a c. 1% sealevel fall at copenhagen due to the associated weakening of gravity within scandinavia (larour et al. 2017). non-land ice processes, including the thermal expansion of seawater, changes in land water storage (i.e. groundwater and dams) and – at the local scale – shifts in atmospheric and oceanic currents, are relatively more important to local sea-level rise at copenhagen (c. 69%) than the global mean (c. 49%). post1993 satellite altimetry indicates that sea level is increasing in the gulf of bothnia, between finland and sweden, more than three times faster than in the north sea (fig. 2). moreover, appreciable local glacio-isostatic adjustment rates mean that global sea-level budget terms cannot be neatly translated into local sea-level budget terms (nielsen et al. 2014). sea-level rise projections while the city-specific sea-level projections we present here are less physically-based than those derived from more complex downscaling approaches (jevrejeva et al. 2016), they are informative within their respective one standard deviation longitude (°e) 5 10 15 20 25 lat itu de (° n ) 52 54 56 58 60 62 64 66 fig. 2 a: trend in mean sea-level elevation (in mm/year) measured by satellite altimetry during the january 1993 and july 2016 period for which data was freely available (nerem et al. 2010). b: present-day (c. 2015) glacioisostatic rebound (in mm/year) simulated by one of the many geodynamic models for which data was freely available (stuhne & peltier 2015). c: relative sea-level change calculated as a minus b, without accounting for geoid differences between both datasets. a longitude (°e) 5 10 15 20 25 lat itu de (° n ) 52 54 56 58 60 62 64 66 n/a < -7 -7 to -5 -5 to -3 -3 to -1 -1 to 1 1 to 3 3 to 5 5 to 7 > 7 b longitude (°e) 5 10 15 20 25 lat itu de (° n ) 52 54 56 58 60 62 64 66 c fig. 2 a: trend in mean sea-level elevation (in mm/year) measured by satellite altimetry during the january 1993 and july 2016 period for which data were freely available (nerem et al. 2010). b: present-day (c. 2015) glacio-isostatic rebound (in mm/year) simulated by one of the many geodynamic models for which data were freely available (stuhne & peltier 2015). c: relative sea-level change calculated as a minus b, without accounting for geoid differences between both datasets. e2019430101-04 uncertainties. the distance between esbjerg and skagen – 280 km – highlights a considerable spatial gradient in sealevel rise. due to ongoing glacio-isostatic rebound at skagen, 21st century sea-level rise relative to 1901–1950 will be limited to 64 ± 28 cm under rcp8.5 and 39 ± 21 cm under rcp4.5 (table 1). at esbjerg, where there is instead ongoing glacio-isostatic subsidence, 21st century sea-level rise will consequently be c. 25 cm greater; 89 ± 28 cm under rcp8.5 and 63 ± 21 cm under rcp4.5. the year 2100 sea-level rise projected for esbjerg under rcp4.5 is therefore similar to that projected for skagen under rcp 8.5. the sea-level forecast for copenhagen is between that of these two endmember case studies. at all three cities, rates of 21st century sea-level rise will be ten times more rapid than rates of 20th century sea-level rise. at esbjerg, the sea-level change over the next century will be approximately equivalent in magnitude to the sea-level change that has occurred there over the past millennium. outlook this study translates the global sea-level projections compiled by the geological survey of denmark and greenland (geus) in support of the 2017 snow, water, ice and permafrost assessment of the arctic monitoring (swipa 2017) and assessment program into local sea-level rise projections at esbjerg, skagen and copenhagen. these danish case studies highlight strong differences in local sea-level histories and projections, as well as marked differences in the drivers of present-day sea-level rise relative to the global average. there is a multi-centennial to millennial lag in the global sea-level response to global climate, which can introduce transient local sea-level responses. contextualising near-term change with long-term perspectives can therefore substantially improve local sea-level projections. contemporary sea-level change is variable across the earth. this study supports ongoing efforts by the department of glaciology and climate at geus to communicate emerging sea-level science to the danish public in a local and regional context, especially with regard to the role of the changing greenland ice sheet (colgan et al. 2018). here, we show that the year 2100 differences in projected sea-level rise between two danish cities under a single climate scenario is approximately equivalent to the differences projected for one city under two climate scenarios. communicating presentday and future sea-level changes throughout the kingdom of denmark – including greenland and the faroe islands – therefore remains a challenging task. references box, j. & colgan, w. 2017: sea level rise contribution from arctic land ice: 1850–2100. in: snow, water, ice and permafrost in the arctic (swipa) 2017, 219–229. arctic monitoring and assessment programme (amap), oslo, norway. church, j. et al. 2013: sea level change. in: climate change 2013: the physical science basis. contribution of working group i to the fig. 3. highlighting the present-day land ice contribution to sea-level rise at copenhagen (a) and the global mean (b) (box & colgan 2017; larour et al. 2017). the slight difference in pie chart size reflects sea-level rise at copenhagen (2.9 ± 0.4 mm/year; holgate et al. 2013) versus the global mean (2.8 ± 0.3 mm/year; nerem et al. 2010) during the 1993–2012 period. non-land ice processes include thermal expansion of seawater, changes in land water storage and shifts in atmospheric and oceanic currents at the local scale. a b russian arctic land ice scandinavian land ice alaskan land ice canadian arctic land ice other land ice antarctic land ice greenland land ice non-land ice processes year skagen copenhagen esbjerg table 1. sea level (m) relative to the 1901–1950 mean at skagen, copenhagen and esbjerg between 0 and 2100 ce based on palaeoreconstructions, tide gauge measurements and projections. uncertainties denote one standard deviation. c. 0 4.0 ± 0.5 n/a –1.5 ± 0.5 1980 –0.01 ± 0.01 0.01 ± 0.01 0.06 ± 0.01 2010 –0.02 ± 0.01 0.02 ± 0.01 0.10 ± 0.01 2040 rcp4.5 0.12 ± 0.08 0.21 ± 0.08 0.28 ± 0.08 rcp8.5 0.15 ± 0.09 0.23 ± 0.09 0.31 ± 0.09 2070 rcp4.5 0.27 ± 0.15 0.37 ± 0.15 0.47 ± 0.15 rcp8.5 0.37 ± 0.18 0.47 ± 0.18 0.57 ± 0.18 2100 rcp4.5 0.39 ± 0.21 0.51 ± 0.21 0.63 ± 0.21 rcp8.5 0.64 ± 0.28 0.77 ± 0.28 0.89 ± 0.28 e2019430101-05 fifth assessment report of the intergovernmental panel on climate change, 1137–1205, cambridge, cambridge university press. https:// doi.org/10.1017/cbo9781107415324.026 colgan, w., grinsted, a., box, j. & macferrin, m. 2018: the mind-bending physics of scandinavian sea-level change. popular science article for sciencenordic.com and videnskab.dk: http://sciencenordic.com/ mind-bending-physics-scandinavian-sea-level-change clemmensen, l., richardt, n. & andersen, c. 2001: holocene sea-level variation and spit development: data from skagen odde, denmark. the holocene 11, 323–331. http://dx.doi. org/10.1191/095968301667877044 gehrels, w., szkomik, k., bartholdy, j., kirby, j., bradley, s., marshall, w., heinemeier, j. & pedersen, j. 2006: late holocene sea-level changes and isostasy in western denmark. quaternary research 66, 288–302. http://dx.doi.org/10.1016/j.yqres.2006.05.004 gehrels, w. & shennan, i. 2015: sea level in time and space: revolutions and inconvenient truths. journal of quaternary science 30, 131–143. http://dx.doi.org/10.1002/jqs.2771 hauerbach, p. 1992: skagen odde – skaw spit. an area of land created between two seas. folia geographica danica 20, 119 pp. holgate, s. et al. 2013: new data systems and products at the permanent service for mean sea level. journal of coastal research 29, 493–504. http://dx.doi.org/10.2112/jcoastres-d-12-00175.1 jevrejeva, s., jackson, l., riva, r., grinsted, a. & moore, j. 2016: coastal sea level rise with warming above 2°c. proceedings of the national academy of sciences 113, 13342–13347. https://doi.org/10.1073/ pnas.1605312113 larour, e., ivins, e. & adhikari. s. 2017: should coastal planners have concern over where land ice is melting? science advances 3, p.e1700537. http://dx.doi.org/10.1126/sciadv.1700537. levermann, a., clark, p., marzeion, b., milne, g., pollard, d., radic, v. & robinson, a. 2013: the multimillennial sea-level commitment of global warming. proceedings of the national academy of sciences 110, 13745–13750. https://doi.org/10.1073/pnas.1219414110 móren, b., sejrup, h., hjestuen, b., borge, m. & schaüble, c. 2018: the last deglaciation of the norwegian channel – geomorphology, stratigraphy and radiocarbon dating. boreas 47, 347–366. http://dx.doi. org/10.1111/bor.12272 nerem, r., chambers, d., choe, c. & mitchum, g. 2010: estimating mean sea level change from the topex and jason altimeter missions. marine geodesy 33, 435–446. http://dx.doi.org/10.1080/0149 0419.2010.491031 nielsen, l., hansen, j., hede, m., clemmensen, l., pejrup, m. & noenygaard, n. 2014: simultaneous estimation of lithospheric uplift rates and absolute sea level change in southwest scandinavia from inversion of sea level data. geophysics journal international 199, 1018–1029. http://dx.doi.org/10.1093/gji/ggu290 permanent service for mean sea level (psmsl) 2018: tide gauge data. retrieved 10 november 2018 from http://www.psmsl.org/data/obtaining. stuhne, g. & peltier, w. 2015: reconciling the ice-6g_c reconstruction of glacial chronology with ice sheet dynamics: the cases of greenland and antarctica. journal of geophysical research: earth surface 120, 1841–1865. http://dx.doi.org/10.1002/2015jf003580 szkornik, k., gehrels, w. & murray, a. 2008: aeolian sand movement and relative sea-level rise in ho bugt, western denmark, during the ‘little ice age’. the holocene 18, 951–965. https://doi. org/10.1177/0959683608091800 how to cite colgan, w., box, j.e., ribeiro, s. & kjeldsen, k.k. 2019: sea-level rise in denmark: bridging local reconstructions and global projections. geological survey of denmark and greenland bulletin 43, e2019430101. https://doi.org/10.34194/geusb-201943-01-01 *corresponding author: william colgan | e-mail: wic@geus.dk 1 geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350, copenhagen k, denmark. https://doi.org/10.1017/cbo9781107415324.026 https://doi.org/10.1017/cbo9781107415324.026 http://sciencenordic.com http://videnskab.dk: http://sciencenordic.com/mind-bending-physics-scandinavian-sea-level-change http://sciencenordic.com/mind-bending-physics-scandinavian-sea-level-change http://dx.doi.org/10.1191/095968301667877044 http://dx.doi.org/10.1191/095968301667877044 https://dx.doi.org/10.1016/j.yqres.2006.05.004 http://dx.doi.org/10.1002/jqs.2771 http://dx.doi.org/10.2112/jcoastres-d-12-00175.1 https://doi.org/10.1073/pnas.1605312113 https://doi.org/10.1073/pnas.1605312113 http://dx.doi.org/10.1126/sciadv.1700537 https://doi.org/10.1073/pnas.1219414110 http://dx.doi.org/10.1111/bor.12272 http://dx.doi.org/10.1111/bor.12272 http://dx.doi.org/10.1080/01490419.2010.491031 http://dx.doi.org/10.1080/01490419.2010.491031 http://dx.doi.org/10.1093/gji/ggu290 http://www.psmsl.org/data/obtaining http://dx.doi.org/10.1002/2015jf003580 https://doi.org/10.1177/0959683608091800 https://doi.org/10.1177/0959683608091800 https://doi.org mailto:wic@geus.dk geological survey of denmark and greenland bulletin 31, 2014, 91-94 91 surface albedo as a proxy for the mass balance of greenland’s terrestrial ice william colgan, jason e. box, robert s. fausto, dirk van as, valentina r. barletta and rene forsberg satellite observations are critical to understanding the mass balance of greenland’s terrestrial ice (fig. 1). the gravity recovery and climate experiment (grace) satellite constellation provides monthly gravimetry observations that can directly assess mass balance. temporal data gaps have begun to appear in the grace record due to declining satellite function. in anticipation of further deterioration in the coverage of grace, we have explored an empirical relation between ice-surface albedo (or reflectance) and ice-mass balance to fill the gaps in the gravimetry record of greenland’s ice-mass balance. as surface albedo observed by the moderate-resolution imaging spectroradiometer (modis) aboard the terra satellite is available in near real-time, employing a modisderived proxy permits near real-time estimates of greenland ice-mass balance. the geological survey of denmark and greenland has begun employing the albedo – mass-balance relation described here to issue near real-time estimates of greenland ice-mass balance during the summer melt season at www.polarportal.org. data and method we employ ice-mass balance data for greenland as assessed by the technical university of denmark monthly grace rl05 solutions from the 2003 to the 2012 summer melt seasons (may to september; barletta et al. 2013). monthly mass balance, calculated by node-centred finite differencing of this solution time series, is available for 44 out of 50 study-period months. single, missing monthly solutions (jun 2003, jun 2011 and jul/aug 2012) prevent resolving mass balance for the months preceding and following the missing solutions (may/jul 2003, may/jul 2011 and jun/sep 2012; tedesco et al. 2013). the 1σ uncertainty associated with monthly mass balance calculated in this way ranges from 45 to 149 gt, with an average of 91 gt over the study period. we take this average value as representative of the uncertainty in gracederived monthly mass balance. we employ the greenland ice-surface albedo observed by terra modis mod10a1 during clear sky conditions. clear sky conditions vary both temporally and spatially, especially in south greenland, where c. 25% of the modis scenes show clear sky. eleven-day running statistics are used to identify and reject <5% of the values within a given scene that exceed 2σ from the running multi-scene mean. to prevent rejecting potentially valid cases, data within 0.04 of the running multi-scene median are not rejected (box et al. 2012). mean monthly albedo is generated from these clear sky and filtered scenes for the melt-season months from 2003 to 2012. we compared the modis monthly albedo with in situ observations from the programme for monitoring of the greenland ice sheet (promice; van as et al. 2013) and the greenland climate network (steffen & box 2001), and found that the root mean squared error reached a minimum of 0.039 in may and a maximum of 0.085 in september. the bias between modis and in situ albedo, which was less than the root mean squared error in all months, ranged from –0.027 in june to 0.022 in september. we therefore take the 1σ uncertainty associated with the ice-sheet-wide modis monthly albedo to be 0.059 during the summer melt season, the average may © 2014 geus. geological survey of denmark and greenland bulletin 31, 91–94. open access: www.geus.dk/publications/bull 50°w 70°n 60°n albedo (no unit) 0.90 0.45 mass balance (kg/m2/a) 150 –1000 65°n 75° a b fig. 1. a: mean annual mass balance of greenland’s terrestrial ice, derived by an inversion of mass loss observed by satellite gravimetry that is constrained by satellite altimetry and fractional ice-coverage information (colgan et al. 2014). b: june to august mean albedo, derived by averaging all available clear sky albedo scenes, each of which has been filtered using running statistics to reject invalid data (box et al. 2012). both datasets span the period from 2004 to 2010 and share a common ice mask where the local ice fraction exceeds 0.5 at 26 km resolution. colour bars saturate at maximum and minimum values. 9292 through september root mean squared error between local in situ and modis observations. a portion of this apparent discrepancy likely results from footprint differences between in situ (c. 10 m2) and modis (c. 500 m2) albedo samples. using these grace-derived mass-balance and modisderived albedo records, we evaluate a single variable regression model to estimate monthly mass balance as a function of monthly albedo. monthly albedo (α) is strongly correlated with monthly mass balance (∂m/∂t) during the summer melt months within the study interval (r = 0.899, p <0.01; fig. 2a). this implies that the greenland ice-mass balance can be statistically approximated by t m a c2 2 a= + where a is a coefficient equal to 1650 gt/month, and c is a constant of –1340 gt/month. a is the apparent sensitivity of mass balance to albedo (e.g. a 0.01 decrease in monthly average albedo corresponds to a 16.5 gt decrease in monthly mass balance), while c would be the theoretical minimum monthly mass balance when all solar radiation is absorbed (e.g. when α = 0). application two sources of error arise when estimating the greenland ice-mass balance via ice-surface albedo: the statistical uncertainty associated with albedo as a proxy for mass balance, and the underlying measurement uncertainty associated with resolving monthly mass balance. the monthly greenland ice-mass balances predicted by single variable albedo regression agree with the monthly mass balances observed by grace within a root mean squared error of ±32 gt/month. combining, in quadrature, this statistical uncertainty with the characteristic measurement uncertainty in the grace-derived greenland ice-mass balance (±91 gt/month), yields a total uncertainty in albedo-regressed mass balance of ±96 gt/month. we calculate uncertainty in, and assess stationarity of, a and c by calculating their values in overlapping four-year subsets of the ten-year study period. this subset analysis yields 1σ uncertainties associated with best-fit a and c parameters of 1650 ±400 and –1340 ±300 gt/month, respectively. an apparent increase in a and decrease in c over time are suggestive of an increase in mass-balance sensitivity to albedo over time. this is consistent with indications that surface mass balance is now the dominant mechanism of greenland ice loss (enderlin et al. 2014). the drift in both a and c, however, is statistically insignificant over the satellite record length presently available for analysis. within associated statistical uncertainty, we therefore suggest that average monthly ice-surface albedo is a stationary proxy for the monthly greenland ice-mass balance during the 2003 to 2012 melt seasons. 0.65 0.7 0.75 0.8 0.85 –350 –300 –250 –200 –150 –100 –50 0 50 100 albedo (no unit) m as s b ala nc e (g t/m on th ) 0.4 0.5 0.6 0.7 0.8 –3000 –2500 –2000 –1500 –1000 –500 0 500 albedo (no unit) m as s b ala nc e (k g/ m 2 / a) y = 1650x – 1340 r = 0.899 n = 44 y = 2230x – 1850 r = 0.694 n = 2615 a b fig. 2. mass balance versus albedo in time and space. a: greenland ice monthly mass balance derived from satellite gravimetry (barletta et al. 2013) versus monthly average albedo derived from satellite imagery (box et al. 2012), during may to september over the period from 2003 to 2012. b: local mean annual mass balance derived from an inversion of satellite gravimetry (colgan et al. 2014) versus local june to august mean albedo derived from satellite imagery (box et al. 2012), averaged over the 2004 to 2010 period and across the domain shown in fig. 1. solid and dashed lines denote ordinary least squares regression (2σ uncertainty envelope). 93 the utility of ice albedo as a proxy for ice-mass balance may be evaluated by comparing graceand modis-derived cumulative mass-balance anomalies and monthly massbalance rates. the albedo-regressed cumulative anomaly captures both the rate and magnitude of mass loss in each melt season between 2003 and 2012 (fig. 3a). albedo-regressed mass-balance rates, however, generally overestimate mass loss early in the melt season (–26 gt/month on average in may and june), and underestimate mass loss late in the melt season (+13 gt/month on average in august and september). the single largest residual is a mass-loss overestimate of –92 gt in june 2009 (fig. 3c). precipitation is recognised to decrease with air temperature, as a function of temperaturedependent absolute humidity. promice weather data suggest the june 2009 outlier is most likely due to anomalously cold air temperatures and little snowfall after the initiation of spring melt, which resulted in an anomalously low june surface albedo and melt rate. while cumulative anomalies compound any systematic biases over the course of a season, the magnitude and associated uncertainty of the albedo regressed, monthly mass-balance rates appear reasonable in the context of analogous grace values (fig. 3b). discussion and summary the mass balance of greenland’s ice reflects a combination of surface mass balance and underlying ice dynamic processes. the physical basis for surface albedo being a skilful proxy of surface mass balance is straightforward; albedo increases with fresh snowfall and decreases with melt or snowpack removal (fig. 2b). albedo therefore integrates the competing surface mass-balance processes of accumulation and ablation. snow or ice albedo directly influences meltwater profig. 4. understanding the relation between surface albedo and mass balance: a schematic overview of previously recognised linkages between increased meltwater runoff and enhanced ice loss in greenland (box & colgan 2013). many intermediate processes convert increased meltwater runoff into increased ice loss via either iceberg discharge or surface or basal mass balance. some processes involving ice-surface albedo and crevasses form positive feedback loops, potentially amplifying mass loss (colgan et al. 2011). a b c –2000 –1500 –1000 –500 0 500 1000 an om aly (g t) grace modis –300 –200 –100 0 100 2003 2005 2007 2009 2011 ra te (g t/m on th ) grace modis 2003 2005 2007 2009 2011 2003 2005 2007 2009 2011 2013 2013 2013 ra te (g t/m on th ) year –100 –50 0 50 100 may june july august september fig. 3. a: cumulative mass balance anomaly for greenland’s ice observed by grace satellite gravimetry over the period from january 2003 to october 2012 (barletta et al. 2013), and the analogous albedo-regressed anomaly for may to september melt periods derived from modis satellite imagery. in each year, cumulative albedo-regressed mass loss is applied to the april anomaly assessed by barletta et al. (2013). b: melt season ice-mass balance rate (barletta et al. 2013), and the analogous albedoregressed rate when grace-derived values are not available. c: residual (modis-derived minus grace-derived) in monthly ice-mass balance during the may to september melt season. climate forcing enhanced process mass loss increased runoff marine convection cryo-hydrologic warming basal lubrication crevasse formation lake drainage submarine ablation terminus instability deformational velocity surface albedo sliding velocity albedo feedback crevasse feedback balance surface mass iceberg discharge balance basal mass 9494 duction and mass loss via runoff (hock 2005). the indirect links between decreased surface albedo (and increased meltwater runoff) and enhanced mass loss via ice discharge from marine-terminating glaciers are numerous and diverse (fig. 4; box & colgan 2013). similar to surface-balance processes, however, processes enhancing ice dynamics, such as crevasses and supraglacial lakes, also generally decrease albedo with increasing mass loss. for example, as crevassed ice absorbs approximately twice as much solar radiation as non-crevassed ice, small changes in crevasse extent can substantially modify albedo. a >10% increase in crevasse extent since c. 1998 within a west greenland study area has been attributed to the acceleration of jakobshavn isbræ (colgan et al. 2011). crevasses can facilitate dynamic mass loss via enhanced terminus instability, as well as enhanced deformational velocity resulting from cryo-hydrologic warming. similarly, a low-albedo ‘dark zone’ forms in the elevation band where meltwater accumulates, both within the snow and firn, as well as in supraglacial lakes. within this ‘dark zone’, up to 40% of variability in annual mass balance is due to variability in summer ice-surface albedo (greuell 2000). in high-melt years, lakes form at higher elevations and have a higher probability of rapidly draining large water volumes to the subglacial hydrological system (liang et al. 2012). albedo ‘dark zone’ width is therefore directly proportional to the delivery of water to the ice–bed interface. given previously recognised strong links between albedo and surface mass balance (hock 2005), the high correlation between greenland ice albedo and mass balance that we have explored supports the notion that the majority of recent greenland ice-mass loss has occurred via meltwater runoff (enderlin et al. 2014). given the numerous and diverse previously postulated links between increased meltwater runoff and enhanced ice-dynamic mass loss, we suggest that a substantial portion of variability in dynamic mass loss is ultimately modulated by surface albedo and meltwater runoff (box & colgan 2013). the overarching inference from our preliminary data exploration is that a strong relation exists between greenland’s ice-surface albedo and mass balance, both through time and across space, and this merits further examination (fig. 2). as near real-time albedo monitoring has proved successful in qualitatively forecasting extreme greenland ice-mass loss events (box et al. 2012), the geological survey of denmark and greenland has begun employing the albedo regression described here to issue near real-time estimates of greenland ice-mass balance during the melt season at www.polarportal.org (fausto et al. 2013). acknowledgement this work is supported by the danish council for independent research, natural sciences (11-115166) and promice. references barletta, v.r., sørensen, l.s. & forsberg, r. 2013: scatter of mass changes estimates at basin scale for greenland and antarctica. the cryosphere 7, 1411–1432. box, j., fettweis, x., stroeve, j.c., tedesco, m., hall, d.k. & steffen, k. 2012: greenland ice sheet albedo feedback: thermodynamics and atmospheric drivers. the cryosphere 6, 821–839. box, j. & colgan, w. 2013: greenland ice sheet mass balance reconstruction. part iii: marine ice loss and total mass balance (1840–2010). journal of climate 26, 6990–7002. colgan, w., steffen, k., mclamb, w.s, abdalati, w., rajaram, h., motyka, r.j., phillips, t. & anderson, r.s. 2011: an increase in crevasse extent, west greenland: hydrologic implications. geophysical research letters 38, l18502. colgan, w., abdalati, w., citterio, m., csatho, b., fettweis, x., luthcke, s., moholdt, g. & stober, m. 2014: hybrid inventory, gravimetry and altimetry (higa) mass balance product for greenland and the canadian arctic. the cryosphere discussions 8, 537–580. enderlin, e.m., howat, i.m., jeong, s., noh, m.-j., van angelen, j.h. & van den broeke, m.r. 2014: an improved mass budget for the greenland ice sheet. geophysical research letters 41, 866–872. fausto, r.s., colgan, w. & langen, p.l. 2013: real-time changes in arctic ice presented in online portal. eos, transactions of the american geophysical union 94, 397–398. greuell, w. 2000: melt-water accumulation on the surface of the greenland ice sheet: effect on albedo and mass balance. geografiska annaler 82a, 489–498. hock, r. 2005: glacier melt: a review of processes and their modelling. progress in physical geography 29, 362–391. liang, y.-l., colgan, w., qin, l, steffen, k., abdalati, w., stroeve, j., gallaher, d. & bayou, n. 2012: a decadal investigation of supraglacial lakes in west greenland using a fully automatic detection and tracking algorithm. remote sensing of environment 123, 127–138. steffen, k. & box, j. 2001: surface climatology of the greenland ice sheet: greenland climate network 1995–1999. journal of geophysical research 106, 33 951–33 964. tedesco, m., fettweis, x., mote, t., wahr, j., alexander, p., box, j.e. & wouters, b. 2013: evidence and analysis of 2012 greenland records from spaceborne observations, a regional climate model and reanalysis data. the cryosphere 7, 615–630. van as, d., fausto, r.s., colgan, w.t., box, j.e. and the promice project team 2013: darkening of the greenland ice sheet due to the meltalbedo feedback observed at promice weather stations. geological survey of denmark and greenland bulletin 28, 69–72. authors’ address es w.c., j.e.b., r.s.f. & d.v.a., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: wic@geus.dk v.r.b. & r.f., technical university of denmark dk-2800 kgs. lyngby, denmark. geological survey of denmark and greenland bulletin 35, 2016, 1-8 1 geological survey of denmark and greenland bulletin 35 • 2016 review of survey activities 2015 edited by adam a. garde, ole bennike, kristine thrane and w. stuart watt geological survey of denmark and greenland ministry of energy, utilities and climate 22 geological survey of denmark and greenland bulletin 35 keywords geological survey of denmark and greenland, survey organisations, current research, denmark, greenland. cover photographs from left to right 1 3d anthroprogenic model of urban infrastructure with buildings and pipes for combination with geological data. 2 elevated plain (c. 800 m a.s.l.) across precambrian basement cut by a deep valley, torngat mountains, labrador, canada. source: google earth. 3 th e geothermal energy potential in denmark is substantial and a webgis portal for exploration of deep geothermal energy based on geolocial and geophysical data has been launched. an interactive 3d tool in the portal gives an intuitive overview of the variations of subsurface topography. 4 aft er drilling two or more overlapping holes in the sea ice, a so-called kajak core with seabed sediments is retrieved for investigations of past climate. photo: jesper hoff mann. frontispiece: facing page field investigations at the sulugssut intrusive complex in the alpine terrain of east greenland. th e rocks formed during the initial stages of the opening of the north atlantic ocean at c. 55 ma. th e fi eld camp in the background was pitched on top of a glacier. photo: th omas f. kokfelt. chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, department of geoscience, aarhus university; minik rosing, geological museum, university of copenhagen; finn surlyk, department of geosciences and natural resource management, university of copenhagen scientifi c editors: adam a. garde, ole bennike, kristine th rane and w. stuart watt editorial secretary: jane holst referees (numbers refer to fi rst page of reviewed article): katrine juhl andresen, dk (27); anonymous (59, 75); per bergmo, no (87); albertas bitinas, lt (47); lars ole boldreel, dk (35); c. kent brooks, uk (59); mikael calner, se (39); andy chadwick, uk (87); jacob q. christensen, dk (9, 47); stefan claesson, se (103); william colgan, ca (71); lynn dafoe, ca (83); christian deibjerg, dk (17); mikael erlström, se (23); ida fabricius, dk (43); wesley farnsworth, no (71); jens galsgaard, dk (31); christopher harrison, ca (63, 83); jens havskov, no (79); michael houmark-nielsen, dk (35); julie hollis, gl (95); jan jeppesen, dk (13); reinhard kirsch, de (23); margrethe kristensen, dk (9); mats larsbo, se (17); gert laursen, dk (13); jerry lloyd, uk (67); anders mattias lundmark, no (103); sebastian mernild, no (75); th orsten nagel, dk (99); allen nutman, au (55); asger ken pedersen, dk (63); toby rivers, ca (91); anders scherstén, se (55); denis schlatter, ch (99); vera schlindwein, de (79); iain sinclair, ca (91); jasna sinigoj, sl (95); jette sørensen, dk (27); svend stouge, dk (39); sander suicmez, dk (43); nicolas r. th ibault, dk (51); clemens ullmann, uk (51); tod waight, dk (31); jacob c. yde, no (71). illustrations: jette halskov, stefan sølberg, susanne rømer, adam a. garde and benny m. schark layout and graphic production: jane holst and annabeth andersen printer: rosendahls-schultz grafi sk a/s, albertslund, denmark manuscripts received: 8 january – 29 february 2016 final versions approved: 11 february – 1 may 2016 printed: 15 july 2016 issn (print) 1604-8156, isbn (print) 978-87-7871-438-1 issn (online) 1904-4666, isbn (online) 978-87-7871-439-8 citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 35, 106 pp. available from geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk and at www.geus.dk/publications/bull © de nationale geologiske undersøgelser for danmark og grønland (geus), 2016 for the full text of the geus copyright clause, please refer to www.geus.dk/publications/bull 3 44 39 middle–upper ordovician and silurian stratigraphy and basin development in southernmost scandinavia n.h. schovsbo, a.t. nielsen and m. erlström 43 types of formation water and produced water in danish oiland gasfi elds: implications for enhanced oil recovery by injection of ‘smart’ water n.h. schovsbo, h.d. holmslykke, c. kjøller, k. hedegaard, l. kristensen, e. th omsen and k.h. esbensen 47 middle pleistocene interglacial deposits near herning, jylland, denmark b.v. odgaard, k.l. knudsen, o. bennike and h.j. granat 51 geochemistry of the maastrichtian rørdal member, jylland, denmark: ce anomaly as a palaeo-redox proxy c. knudsen and b.w. lauridsen greenland canada kenya ethiopia nigeria cameroon zambia ghana mozambique tanzania denmark norway iceland sweden finland 7 review of survey activities 2015 f.g. christiansen 9 miocene deposits at silkeborg, jylland, and their infl uence on hydrology p.r jakobsen, e.s. rasmussen, k. dybkjær and j. kidmose 13 3d hydrogeological modelling for urban subsurface management in odense, denmark s. mielby, t.m. pallesen and p.b.e. sandersen 17 monitoring of pesticide leaching from cultivated fi elds in denmark w. brüsch, a.e. rosenbom, n. badawi and p. olsen 23 a webgis portal for exploration of deep geothermal energy based on geological and geophysical data h. vosgerau, a. mathiesen, m.s. andersen, l.o. boldreel, m.l. hjuler, e. kamla, l. kristensen, c.b. pedersen, b. pjetursson and l.h. nielsen 27 towards a national 3d geological model of denmark p.b.e. sandersen, t. vangkilde-pedersen, f. jørgensen, r. th omsen, j. tulstrup and j. fredericia 31 pre-quaternary rocks and sediments with a high level of radioactivity in denmark p. gravesen and p.r. jakobsen 35 tectonic control on the formation of roskilde fjord, central sjælland, denmark s.a.s. pedersen and p. gravesen 5 grey indicates countries where geus has ongoing or recently completed projects. orange indicates countries with geus projects described in this volume. 55 new zircon u-pb and hf isotopic constraints on the crustal evolution of the skjoldungen region, southeast greenland t.f. kokfelt, t. næraa, k. th rane and l. bagas 59 in situ fractionation and inward migration of the solidifi cation front in the skaergaard intrusion, east greenland t.f.d. nielsen 63 palaeovalleys at the basal unconformity of the palaeoproterozoic karrat group, west greenland p. guarnieri, c.a. partin and d. rosa 67 investigations of past climate and sea-ice variability in the fj ord area by station nord, eastern north greenland n. nørgaard-pedersen, s. ribeiro, n. mikkelsen, a. limoges and m.-s. seidenkrantz 71 placing greenland ice sheet ablation measurements in a multi-decadal context d. van as, r.s. fausto, j. cappelen, r.s.w. van de wal, r.j. braithwaite, h. machguth and the promice project team 75 regional climate-model performance in greenland fi rn derived from in situ observations c. charalampidis, d. van as, p.l. langen, r.s. fausto, b. vandecrux and j.e. box 79 crustal structure over the nagssugtoqidian deformation front in west greenland: receiver function analysis t. dahl-jensen, p.h. voss and t.b. larsen 83 new geophysical and geological mapping of the eastern baffi n bay region, off shore west greenland u. gregersen, p.c. knutz and j.r. hopper 87 mapping of the co2 storage potential in the nordic region k.l. anthonsen, p. frykman and c.m. nielsen 91 burial and exhumation history of the labradornewfoundland margin: fi rst observations p. japsen, p.f. green, j.m. bonow, a.m. hinchey and d.h.c. wilton 95 th e greenland mineral resources portal – another step forward m. pedersen, m. hansen, b.h. heincke and l. th orning 99 afieldwork – an android app for offl ine recording of geological information and data display m. hansen, m.n. petersen, t.f. kokfelt and b.m. stensgaard jagedisplay: soft ware for evaluation of data distributions in u-th -pb geochronology t.b. th omsen, t. heijboer and p. guarnieri vietnam tajikistan 103 66 7 review of survey activities 2015 flemming g. christiansen deputy director 2015 was a tough year for many geologists around the world, and the years to come may be even tougher. low prices of oil, gas and other energy sources, and also of many mineral commodities have led to a significantly lower level of investments in exploration and production than seen in previous years. both society and industry suffer from much lower income, so cost reduction is the new buzzword; and many geologists have lost their jobs ‒ also in denmark and at geus. investments in data, projects and knowledge ‒ including many typical products from the geological survey of denmark and greenland (geus) ‒ are worryingly low and will eventually reduce the level of knowledge for decision makers in both the private and public sectors. commodity prices are, however, cyclic by nature, and it is very important for geus to continue to collect and compile new data, build up new knowledge and models based on research projects – and prepare for a future that will still rely on traditional resources but with a strong focus on a transition towards new green technologies that can reduce the consumption of fossil fuels and the emission of co 2 . this issue of geus’ review of survey activities has a broader content and a total of 24 four-page articles, more than seen for several years. it reflects that geus works on many different subjects and is preparing for a future where geology still plays an important part in the planning of important activities in denmark and greenland. eleven papers are on denmark, eight on greenland and five on international and general themes. activities in denmark geus’ danish activities and research cover a wide range of topics within our specific programme areas: data, water, energy, mineral resources and nature and climate, as well as many other basic research projects. the use of groundwater is very important for denmark, and geus carries out many studies on water resources, their protection and possible future challenges due to changes in climate and use. as a follow-up on previous studies of the miocene sedimentary succession in jylland, one paper focuses on deposits at silkeborg and their influence on hydrology. another paper describes 3d hydrogeological modelling important for the urban subsurface management in odense. a third paper is on the monitoring of pesticide leaching from cultivated fields in denmark, an activity that has been carried out for several decades. denmark has a large potential for subsurface geothermal energy. following several regional studies and many local case stories in preparation for drilling campaigns, geus has developed a webgis portal for exploration of deep geothermal energy based on geological and geophysical data. one paper describes this portal where the relevant geological and geophysical maps, data and key information from drilling are easily accessed. many of the applied studies emphasise the strong need for a comprehensive national 3d geological model of denmark. similar models are being developed in several other european countries in order to provide the best possible background for large resource and infrastructure projects. one paper describes the strategy behind the effort that geus is currently putting into developing such models and gives examples of how existing data and geological models can be applied. sediments and rocks in denmark generally have a low content of radioactive minerals and radon. some of the highest levels of radioactivity on mors and bornholm have been studied in detail to understand their geological control and distribution in space. this is described in one paper. another paper describes how the terrain in the roskilde area, and in particular the roskilde fjord, is controlled by deeper tectonic features that are clearly recorded in subsurface maps of the danian limestone and paleocene chalk, marl and clay. after several studies of the palaeozoic succession in denmark, it is now possible to make a complete middle‒upper ordovician and silurian stratigraphy and basin model for all of southern scandinavia. this is described in another paper using wireline logs as a correlation tool. several methods of enhanced oil recovery, e.g. injection of ‘smart’ water, have the potential to significantly increase oil production, and thereby income for danish society. to do this, a detailed understanding of different types of reservoirs is required, as shown in a paper describing the distribution and composition of primary formation water and produced water in the north sea oilfields. © 2016 geus. geological survey of denmark and greenland bulletin 35, 7–8. open access: www.geus.dk/publications/bull 88 interglacial marine deposits are fairly common in denmark but for the first time such deposits are described from kibæk in the herning area, where they occur several metres above present sea level. chalk is a very important rock for denmark. it is the reservoir for most of the oil and gas resources in the north sea, it is important as a groundwater reservoir in large areas onshore, and it is an important resource for cement production. one paper provides new details of geochemical stratigraphy as a tool to understand the depositional environments and their geographical correlation. activities in greenland once again there was a high level of activity in and about greenland in 2015. many large and small projects were carried out, studies that are important for evaluating and marketing the resource potential in greenland. the level of industrial activity in both oil and mineral exploration is very low at the moment, but it is important to prepare for a future when prices of the most important commodities will eventually rise again. for the last few years geus has focused on mapping and mineral evaluation activities in south-east greenland; this emphasis is now changing to west and north-west greenland. two papers give new information from southeast greenland. one paper describes new geochronological data using hf isotopes in zircon to unravel archaean crustal accretion processes in the skjoldungen region, and another provides entirely new insight into the fractionation processes of the world-famous skaergaard intrusion, which has been discussed for many decades. the last paper on greenland bedrock geology describes palaeoproterozoic palaeovalleys underlying the karrat group in the uummannaq region of west greenland, where new field activity has started. studies and monitoring of the greenland ice sheet and studies of marine sediments in the fjords around greenland result in important contributions from geus to global climate models. one paper investigates past climate and sea-ice variability in the remote eastern part of north greenland close to station nord. results from the important promice monitoring project with its focus on temperature sensitivity of ice sheet-ablation are presented in another paper. a third paper is on climate models based on in situ observations of greenland firn. passive, so-called receiver function analysis of seismological signals from natural earthquakes around the world can be used to information on the deep crustal structure, e.g. the depth to moho. one paper presents results from such a study across the nagssugtoqidian front in west greenland where a significant change in moho is recorded across an ancient plate boundary and subduction zone. systematic mapping of offshore sedimentary deposits and evaluation of the petroleum potential is crucial for planning and marketing of licensing rounds and for advice on subsequent exploration activities. after several successful licensing rounds and a relatively high level of activity, the seismic data coverage in baffin bay is now relatively high compared to many other regions around greenland. based on a systematic updated interpretation and mapping, one paper outlines some interesting possibilities for future exploration but also describes the main risks and uncertainties. international and broader technical themes geus also works overseas in many different countries with a variety of projects and is involved in broader thematic studies. over the years geus has been involved in several projects of carbon dioxide capture and storage (ccs), especially with european union and industrial funding. one paper summarises mapping of the storage capacity of sandstone aquifers in denmark and the nordic region, based on a large nordic collaboration project. another paper describes a preliminary study of the burial and exhumation history of the labrador-newfoundland margin using apatite fission track analysis and thermal maturity methods similar to previo us studies on the greenland side of the labrador sea. geus is constantly developing databases and facilities for easier access to and use of data, as well as new methodologies and technologies to make research easier, quicker and better for our geologists. one paper describes the greenland mineral resources portal; a new version of this was launched by geus and the ministry of mineral resources in greenland (mmr) in april 2015. the portal is based on decades of work and re-organisation of previous databases with an ambition of securing data and making them easily accessible through the internet. another paper is on a new app – afieldwork – that has been designed specifically for field work in greenland to make digital capture of data as easy and efficient as possible. the last paper presents the development of special software for evaluation of data distribution in u-th-pb geochronology – jagedisplay. geological survey of denmark and greenland bulletin 1, 948 pp. geological survey of denmark and greenland bulletin 1 • 2003 the jurassic of denmark and greenland edited by jon r. ineson and finn surlyk geological survey of denmark and greenland ministry of the environment geological survey of denmark and greenland bulletin 1 keywords northwest europe, denmark, greenland, sweden, the netherlands, jurassic, chronostratigraphy, biostratigraphy, lithostratigraphy, sequence stratigraphy, structural history, basin evolution, sedimentology, palynostratigraphy, geochemistry, coal petrography/palynology cover palaeogeography of the mesozoic rift system in the north atlantic region in the middle jurassic (c. 160–180 ma), viewed towards the north. reconstruction by stefan sølberg, based on the palaeogeographic maps of ziegler (1990) and doré (1992). for references, see surlyk (2003, this volume). chief editor of this series: peter r. dawes scientific editors: jon r. ineson and finn surlyk copy editors: jon r. ineson and birgit eriksen editorial secretary: birgit eriksen critical readers: see list on page 6 illustrations (geus): stefan sølberg, gurli e. hansen bengaard, jette halskov, eva melskens, helle zetterwall photographic work (geus): jacob lautrup, benny m. schark, peter k. warna-moors lay-out and graphic production: carsten e. thuesen printers: schultz grafisk, albertslund, denmark manuscripts submission/acceptance dates: see individual articles printed: 28th october 2003 isbn 87-7871-116-9 geological survey of denmark and greenland bulletin the series geological survey of denmark and greenland bulletin replaces geology of denmark survey bulletin and geology of greenland survey bulletin. citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary the following form is suggested: geol. surv. den. green. bull. 1, 948 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk or geografforlaget aps fruerhøjvej 43, dk-5464 brenderup, denmark phone: +45 63 44 16 83, fax: +45 63 44 16 97, e-mail: go@geografforlaget.dk  danmarks og grønlands geologiske undersøgelse (geus), 2003 5 6 7 9 23 61 75 115 147 217 233 247 265 301 349 367 403 contents dedication referees preface the jurassic of denmark and greenland: key elements in the reconstruction of the north atlantic jurassic rift system f. surlyk and j.r. ineson stratigraphy the lower jurassic of europe: its subdivision and correlation k.n. page the middle jurassic of western and northern europe: its subdivisions, geochronology and correlations j.h. callomon the upper jurassic of europe: its subdivision and correlation a. zeiss the jurassic dinoflagellate cyst zonation of subboreal northwest europe n.e. poulsen and j.b. riding denmark, southern sweden and the netherlands jurassic lithostratigraphy and stratigraphic development onshore and offshore denmark o. michelsen, l.h. nielsen, p.n. johannessen, j. andsbjerg and f. surlyk the jurassic of the netherlands g.f.w. herngreen, w.f.p. kouwe and th.e. wong danish central graben upper jurassic – lower cretaceous of the danish central graben: structural framework and nomenclature p. japsen, p. britze and c. andersen middle jurassic – early cretaceous rifting of the danish central graben j.j. møller and e.s. rasmussen sequence stratigraphy of the jurassic of the danish central graben j. andsbjerg and k. dybkjær sedimentology and sequence stratigraphy of the bryne and lulu formations, middle jurassic, northern danish central graben j. andsbjerg the use of spectral natural gamma-ray analysis in reservoir evaluation of siliciclastic sediments: a case study from the middle jurassic of the harald field, danish central graben i.l. fabricius, l. dahlerup fazladic, a. steinholm and u. korsbech sedimentology and sequence stratigraphy of paralic and shallow marine upper jurassic sandstones in the northern danish central graben p.n. johannessen volgian–ryazanian ‘hot shales’ of the bo member (farsund formation) in the danish central graben, north sea: stratigraphy, facies and geochemistry j.r. ineson, j.a. bojesen-koefoed, k. dybkjær and l.h. nielsen 3 439 459 527 543 555 585 611 631 659 723 777 813 865 893 931 danish basin and fennoscandian border zone triassic and jurassic transtension along part of the sorgenfrei–tornquist zone in the danish kattegat t.e. mogensen and j.a. korstgård late triassic – jurassic development of the danish basin and the fennoscandian border zone, southern scandinavia l.h. nielsen the jurassic of skåne, southern sweden a. ahlberg, u. sivhed and m. erlström an offshore transgressive–regressive mudstone-dominated succession from the sinemurian of skåne, sweden n. frandsen and f. surlyk lower jurassic (pliensbachian) ammonites from bornholm, baltic sea, denmark d.t. donovan and f. surlyk the lower–middle jurassic of the anholt borehole: implications for the geological evolution of the eastern margin of the danish basin o.b. nielsen, m.-s. seidenkrantz, n. abrahamsen, b.j. schmidt, e.b. koppelhus, h. ravn-sørensen, u. korsbech and k.g. nielsen burial depth and post-early cretaceous uplift of lower–middle jurassic strata in the fennoscandian border zone based on organic maturity h.i. petersen, l.h. nielsen, t. bidstrup and e. thomsen early and middle jurassic mires of bornholm and the fennoscandian border zone: a comparison of depositional environments and vegetation h.i. petersen, l.h. nielsen, e.b. koppelhus and h.s. sørensen east greenland the jurassic of east greenland: a sedimentary record of thermal subsidence, onset and culmination of rifting f. surlyk palynostratigraphy and palaeoenvironments of the rævekløft, gule horn and ostreaelv formations (lower–middle jurassic), neill klinter group, jameson land, east greenland e.b. koppelhus and g. dam palynostratigraphy and palaeoenvironment of the middle jurassic sortehat formation (neill klinter group), jameson land, east greenland e.b. koppelhus and c.f. hansen shallow marine syn-rift sedimentation: middle jurassic pelion formation, jameson land, east greenland m. engkilde and f. surlyk the jurassic of kuhn ø, north-east greenland p.c. alsgaard, v.l. felt, h. vosgerau and f. surlyk stratigraphy and sedimentology of a basement-onlapping shallow marine sandstone succession, the charcot bugt formation, middle–upper jurassic, east greenland m. larsen, s. piasecki and f. surlyk shelf-edge delta and slope deposition in the upper callovian – middle oxfordian olympen formation, east greenland m. larsen and f. surlyk 4 5 this book is dedicated to the memory of ole winther christensen (1951–1998), director of the geological survey of denmark (dgu) from 1990 to 1995 and the geological survey of denmark and greenland (geus) from 1995 to 1998. from its conception in the early 1990s, the ‘jurassic book’ has benefited from the progressive integration of danish geological institutions. this began in 1995 with the amalgamation of dgu with the geological survey of greenland (ggu) to form geus and culminated in 2002 with the opening of the geocenter copenhagen, a conglomeration of geus, the danish lithosphere centre (dlc), the geological museum and the geological and geographical institutes of the university of copenhagen. ole winther christensen played a key role in both these positive developments in danish geology. sadly, he was not to experience the full realisation of the geocenter copenhagen concept due to his untimely death in 1998. dedication 6 j. alexander, university of east anglia, uk j. andsbjerg, geological survey of denmark and greenland, denmark d.j. batten, university of wales, uk d.g. benson, texas, usa j.h. callomon, university college london, uk j. cartwright, cardiff university, uk b.m. cox, british geological survey, uk r.j. davey, simon petroleum technology ltd, uk r.j. davies, mobil north sea ltd, uk g. dietl, staatliches museum für naturkunde, germany d.t. donovan, university college london, uk t. dreyer, norsk hydro research centre, norway k. dybkjær, geological survey of denmark and greenland, denmark c.j. fielding, university of queensland, australia m.j. fisher, helensburgh, uk j. gjelberg, norsk hydro research centre, norway m.b. gowers, dong norge as, norway f. gramann, niedersächsisches landesamt für bodenforschung, germany d. guy-ohlsen, swedish museum of natural history, sweden a. hallam, birmingham university, uk s.d. harker, totalfinaelf exploration, uk g.f.w. herngreen, tno-nitg, the netherlands s.p. hesselbo, university of oxford, uk s. holloway, british geological survey, uk a. hurst, university of aberdeen, uk l.n. jensen, statoil norge as, norway j.a. korstgård, university of aarhus, denmark d.a. leckie, geological survey of canada, canada h.b. lindgreen, geological survey of denmark and greenland, denmark c. mangold, université claude bernard lyon i, france j.e.a. marshall, university of southampton, uk l.h. nielsen, geological survey of denmark and greenland, denmark a. nøttvedt, norsk hydro canada, canada s. olaussen, norsk agip as, norway t. olsen, statoil norge as, norway k.n. page, university of plymouth, uk n. parkinson, western atlas logging services, uk g.k. pedersen, university of copenhagen, denmark r.m. pegrum, statoil norge as, norway s. piasecki, geological survey of denmark and greenland, denmark k.t. pickering, university college london, uk d. pirrie, university of exeter, uk a.g. plint, university of western ontario, canada j.b. riding, british geological survey, uk a. ryseth, norsk hydro research centre, norway b.w. sellwood, reading university, uk r.j. steel, university of wyoming, usa g.m. sykes, helix rds ltd, uk n.h. trewin, university of aberdeen, uk r.v. tyson, university of newcastle, uk g. warrington, british geological survey, uk o.v. vejbæk, geological survey of denmark and greenland, denmark a.g. whitham, casp, uk a. wierzbowski, warsaw university, poland p.b. wignall, leeds university, uk referees the editors are indebted to the following referees, whose conscientious and incisive reviews have been invaluable in the production of this book – your efforts are greatly appreciated. the jurassic sedimentary successions of denmark and east greenland accumulated on opposite sides of a complex rifted seaway between present-day greenland and northwest europe. the mesozoic–cenozoic sedimentary basins created along this seaway are of major importance both scientifically, as they preserve a record of the early evolution of the north atlantic region, and economically as one of the rift arms contains the north sea petroleum province. the jurassic system, in particular, has been the focus of intensive study in northwest europe and greenland. not only has this system, since the days of william smith, been at the forefront of stratigraphic research but it also forms a critical component of the north sea hydrocarbon province, yielding both the most important source rocks and a wide range of sandstone reservoirs. although the stratigraphic development of the jurassic in denmark and east greenland can be compared at a number of levels, the nature of the occurrences in the two regions is very different. the jurassic of east greenland is one of the world’s best-exposed ancient rift basins and is widely regarded as a classic ‘field laboratory’. the jurassic strata are exposed in spectacular cliff sections that provide unique opportunities for detailed research into process sedimentology, genetic stratigraphy and 3d sedimentary architecture. the danish jurassic strata, in contrast, have limited outcrop but are well known from the subsurface, both on land and beneath the waters of the north sea. the papers collected in this volume reflect this contrast – the stratigraphic evolution of east greenland has been deciphered primarily on the basis of detailed outcrop geological studies whereas the corresponding stratigraphic analyses of the danish basin and the danish sector of the central graben are largely dependent on ‘remote’ subsurface data. jurassic stratigraphic research in denmark over the last two decades has benefited immensely from the interaction between these two contrasting yet complementary approaches. the origins of this book go back to the early 1990s when the idea was mooted for a book on the ‘jurassic of denmark and adjacent areas’, initially with a view to publication of the main results of ph.d. studies that were underway at the geological survey of denmark (dgu) at that time. in 1995, with the amalgamation of dgu with the geological survey of greenland (ggu) to form the geological survey of denmark and greenland (geus), the conceptual framework of the book expanded to include the jurassic of east greenland, a research area that was under sharp focus both at ggu and at the university of copenhagen. as the editing of the book entered the final phase, the geological survey relocated to the new geocenter copenhagen – a centralised amalgam of the survey (including the danish lithosphere centre) and the geological and geographical institutes and the geological museum of the university of copenhagen. from conception to publication, therefore, the book charts the changing structure of some of the central geological research bodies in denmark, and its completion coincided with the inception of a new integrated natural science research centre. the central aim of the book is to present the results of an intense period of research activity in denmark on the jurassic system over the last fifteen years – and, where relevant, to present these results at a comprehensive level that is almost impossible in modern scientific journals. although covering a range of subjects, the common thread that runs through the book is the detailed documentation of the history of the jurassic rift system as recorded in the sedimentary basins of greenland and denmark. particular areas of focus include: (1) the sedimentary and stratigraphic signatures of syn-rift successions, whether revealed by detailed outcrop study or on the basis of integrated reflection seismic, petrophysical and core data; and (2) testing and application of sequence stratigraphic models and concepts at a variety of scales and in different structural settings. although focussing on broad geoscientific topics of general relevance, the book also provides data of specific value to the hydrocarbon industry. the danish basin and, in particular, the danish central graben are prospective basins with exploration histories stretching back nearly fifty years. a number of jurassic fields are under development and production in the danish central graben, and exploration interest remains high. the structural, sedimentological and stratigraphic papers in this volume thus represent a direct source of essential data for the hydrocarbon industry. the onshore east greenland basins, in contrast, are not prospective per 7 preface se, yet the detailed sedimentological and stratigraphic analyses included here will be of particular interest to petroleum geologists both as direct stratigraphic analogues of the succession on the conjugate margin (midnorway shelf) and as reservoir analogues or case studies applicable particularly to the north sea region but also valid elsewhere. introductory chronostratigraphic reviews of the lower, middle and upper jurassic were planned from the outset, and contributions were solicited from three international authorities in this field, together with a paper on the jurassic of southern sweden. furthermore, a review of the jurassic of the netherlands was invited from the geological survey of the netherlands (rgd) for comparative purposes, building on previous close stratigraphic co-operation between dgu and rgd in the late 1980s. the aim has been to produce a book that is as balanced and consistent as possible, in terms of content, terminology and appearance. given the range of subjects covered, however, a certain degree of heterogeneity is inevitable and full consistency in terminology cannot be achieved. the gradstein et al. (1994) timescale is used in most cases but the haq et al. (1988) and harland et al. (1990) time-scales are employed by some authors; in all cases, the origin of the time-scale used is clearly indicated. several forms of chronostratigraphic terminology are in common use, all being inherently logical and fully acceptable; particularly prevalent are the ‘standard zone’ nomenclature (callomon & donovan 1974) and the ‘chronozone’ terminology, as laid down in the international stratigraphic guide (salvador 1994). editorial flexibility has been exercised here, although consistency within individual articles was required. to enhance uniformity, a common graphical style has been imposed wherever possible; detailed sedimentary logs are somewhat variable, however, being dictated by different individual styles and demands. in an enterprise of this type, undertaken over a number of years, there are clearly many people both in denmark and abroad who have helped us towards publication. the research projects that formed the initial stimulus behind the book were supported both by state funding – the danish energy agency (energy research program, efp), the danish natural science research council (snf), the danish research academy and the norwegian petroleum directorate (npd) – and by the private sector, including amerada hess, amoco, british petroleum, the carlsberg foundation, conoco, mærsk olie og gas, norsk hydro, saga petroleum and statoil. the long-term support of danish geological research by these funding bodies and companies is gratefully acknowledged. we are also indebted to a long list of international referees; their contribution is acknowledged elsewhere but their importance in upholding the international standard of the papers bears repetition. during the scientific and technical editing phase, we have leaned heavily on three key personnel: hanne b. sørensen, who converted editorial hieroglyphics into ordered manuscripts; birgit eriksen, who meticulously checked final manuscripts and proof copies; and stefan sølberg whose skilled graphical imprint is engraved on almost every illustration in the book. on editorial matters, we have also benefited greatly from close co-operation with peter r. dawes and esben w. glendal in the editorial office at geus. in the latter stages we have been increasingly reliant on the professional layout work by carsten e. thuesen. to all the above, we offer our heartfelt thanks. jon r. ineson finn surlyk on behalf of the ‘jurassic book’ convening group: jon r. ineson, finn surlyk, karen dybkjær, lars. h. nielsen, niels e. poulsen. references callomon, j.h. & donovan, d.t. 1974: a code of mesozoic stratigraphic nomenclature. in: colloque du jurassique à luxembourg 1967. mémoire du bureau de recherches géologiques et minières 75, 75–81. gradstein, f.m., agterberg, f.p., ogg, j.g., hardenbol, j., van veen, p., thierry, j. & huang, z. 1994: a mesozoic time scale. journal of geophysical research 99, 24051–24074. haq, b.u., hardenbol, j. & vail, p. 1988: mesozoic and cenozoic chronostratigraphy and cycles of sea-level change. in: wilgus, c.k. et al. (eds): sea-level changes – an integrated approach. society of economic paleontologists and mineralogists special publication 42, 71–108. harland, w.b., armstrong, r.l., cox, a.v., craig, l.e., smith, a.g. & smith, d.g. 1990: a geologic time scale 1989, 263 pp. cambridge: cambridge university press. salvador, a. (ed.) 1994: international stratigraphic guide. a guide to stratigraphic classification, terminology, and procedure, 2nd edition, 214 pp. boulder, colorado: international union of geological sciences and geological society of america, inc. 8 geological survey of denmark and greenland bulletin 41, 2018, 39-42 39 geological maps are core products of national geological surveys and represent the sum of geological knowledge of any given area. however, dedicated and extensive mapping projects in the arctic are mostly a thing of the past due to difficulty in financing such costly basic research efforts. today, an overview of the geology of greenland is portrayed by a seamless digital 1:500  000 scale geological map (kokfelt et al. 2013; pedersen et al. 2013), based on printed maps on this scale produced since 1982 by the geological survey of denmark and greenland (geus; see holst et al. 2013). the digital map now makes it possible to update smaller areas with new, published or otherwise quality-controlled geological data (e.g. kolb et al. 2016). this ensures that the map reflects the current state of geological knowledge without undertaking extensive new mapping to update individual map sheets, as has previously been the modus operandi. an online version of the map is available from www.greenmin.dk/map. however, procedures are required to ensure that updates are carried out routinely and that the quality and coherence of the updated map is of the survey’s standards. results of recent field work in the wandel sea basin (fig. 1) and in particular the publication of a new geological map sheet kilen on a scale of 1:100 000 (svennevig in press) have implications for the geology shown on the above mentioned 1:500 000 scale seamless geological map of greenland. the post-devonian part of this map in eastern north greenland has been updated according to the results of studies published since the publication of the original printed maps (bengaard & henriksen 1986; jepsen 2000). the changes do not call for an update of the 1:2 500 000 scale geological map of greenland (henriksen et al. 2009). update of the seamless 1:500 000 scale geological map of greenland based on recent field work in the wandel sea basin, eastern north greenland kristian svennevig greenland triassic older rocks faults carboniferous–permian palaeogene jurassic–cretaceous dykes, mostly late cretaceous quaternary 36°w 28°w 20°w 83°n 12°w 82°n 81°n hffz tlfz peary land eg fz holm land amdrup land kilen wandel hav kap washington 100 km kim fjelde kctz gr øn la nd sh av et kronprins christian land inland ice kap rigsdagen station nord prinsesse thyra ø fig. 1. simplified geological map of the wandel sea basin showing the major structural elements and upper palaeozoic to palaeogene sediments. modified from bengaard & henriksen (1986) and henriksen (2003) with updates from kilen modified from svennevig (in press). egfz: east greenland fault zone. hffz: harder fjord fault zone. kctz: kap cannon thrust zone. tlfz: trolle land fault zone. red frame shows the position of fig. 2 and black arrow shows the new occurrence of thyra ø formation. © 2018 geus. geological survey of denmark and greenland bulletin 41, 39–42. open access: www.geus.dk/bulletin http://www.greenmin.dk/map http://www.geus.dk/bulletin 4040 updates to the geological map based on recent field work in the wandel sea basin parts of the seamless 1:500 000 geological map of greenland (kokfelt et al. 2013; pedersen et al. 2013) have been updated. these updates are based on data from extensive field work in 2012, 2013 and 2016 in the carboniferous–palaeogene wandel sea basin in eastern north greenland (bojesenkoefoed et al. 2014; hovikoski et al. in press; piasecki et al. in press), the results of a recent phd-thesis (svennevig in press; svennevig et al. 2016, 2017) and previously published work (pedersen & håkansson 1999; henriksen 2003). in the course of this work, two old map units have been emended and four new ones established. furthermore, four map units have been expanded to encompass new areas. the legend of the seamless map has been updated accordingly. twenty-one lithological units from the new 1:100  000 scale map sheet (svennevig in press) are reduced to five units in the 1:500 000 scale map, and the revised stratigraphic resolution is seen in fig. 2b. the letter codes mentioned below (if, c2, mm, lp, mi, tr, j, jc, lc and uc) refer to the code used in the geus map database. 1. expansion of unit c2: upper carboniferous sediments. a small occurrence of tentatively dated upper palaeozoic melange unit at kilen is included in the unit ‘upper carboniferous sediments’ (c2) found on amdrup and holm land. 25 km g r ø n l a n d s h a v e t f l a d e i s b l i n k 25 km g r ø n l a n d s h a v e t f l a d e i s b l i n k i ii gt cd st normal fault upper palaeozoic older rocks carboniferous palaeogene upper cretaceous jurrasic upper permian syncline anticlinejurrasic cretaceous quaternary generic fault reverse fault normal fault detachment fault jurassic triassic older rocks carboniferous lower permian upper permianpalaeogene upper cretaceous lower cretaceous syncline anticline quaternary generic fault a b nakkehovednakkehoved station nord prinsesse ingeborg halvø prinsesse ingeborg halvø erik s. henius land nordostrundingennordostrundingen k r o n p r i n s c h r i s t i a n l a n d k r o n p r i n s c h r i s t i a n l a n d kilen kilen amdrup landamdrup land 81°n 81°n station nord erik s. henius land wandel hav wandel hav 15°w 15°w fig. 2. geological maps of kronprins christian land showing the geology before (a) and after (b) the update of the digital map. one of the items on the new map (fig. 2b) is a n–s-verging fold and thrust belt covering prinsesse ingeborg halvø and kilen as described by svennevig (in press) and svennevig et al. (2016, 2017). faults on fig. 2b adopted from svennevig (in press): i and ii: normal fault i and ii. cd: central detachment, st: sadelfjeld thrust. gt: gåseslette thrust. 41 furthermore, the ‘mallemuk mountain group’ (mm) in kim fjelde on peary land, a relic from the old printed map (bengaard & henriksen 1986), is included in the c2 unit following (stemmerik et al. 1996, 2000) and the map unit ‘mallemuk mountain group’ (mm) has thus been emended. the unit is also found on prinsesse ingeborg halvø (see below, fig. 2). 2. erection of a new unit tr: triassic siliciclastic sediments, undivided. the newly discovered triassic isrand formation at kilen and the triassic on amdrup land (alsen et al. 2017), is combined in a new map unit called ‘triassic siliciclastic sediments, undivided’ (tr). the strata on amdrup land were previously mapped as jurassic (fig. 2a; stemmerik et al. 2000). 3. erection of a new unit j: jurassic siliciclastic sediments, undivided. jurassic formations at kilen: the gletscherport, mågensfjeld, birkelund fjeld and kuglelejet formations (hovikoski et al. in press; svennevig et al. in press) are combined into the new map unit ‘jurassic siliciclastic sediments, undivided’ (j) which is only found at kilen. 4. erection of a new unit lc: lower cretaceous siliciclastic sediments, undivided. the lowermost lower cretaceous formations at kilen, the dromledome and lichenryg formations, are combined with the upper aptian – early to middle cenomanian galadriel fjeld formation (hovikoski et al. in press; svennevig et al. in press) in the new map unit ‘lower cretaceous siliciclastic sediments, undivided’ (lc). 5. expansion of the unit uc: upper cretaceous siliciclastic sediments, undivided. the sølverbæk formation at kilen (hovikoski et al. in press; svennevig et al. in press) is included in the unit uc, ‘upper cretaceous siliciclastic sediments, undivided’, together with small occurrences of upper cretaceous sediments at kap washington in northernmost greenland. a number of regionally significant faults and fold axes from the 1:100  000 scale kilen map sheet have also been adapted for the 1:500 000 scale map. these are the normal fault i and ii and the reverse faults of the central detachment, the gåseslette thrust, and the saddelfjeld thrust (fig 2b; svennevig in press; svennevig et al. in press). updates based on published work on prinsesse ingeborg halvø the geology on prinsesse ingeborg halvø has also been updated (fig. 2). in the original version of the 1:500 000 scale map this area was mapped as mainly “undivided upper proterozoic to lower palaeozoic sediments” (bengaard & henriksen 1986; kokfelt et al. 2013). since the publication of the printed map in 1986, a map was published as fig. 7 in pedersen & håkansson (1999). this figure was also adapted for a 1:1 000 000 scale map of the caledonian orogeny (henriksen 2003). the 1:500  000 scale seamless map has been updated according to this and now contains three carboniferous–permian map units of the wandel sea basin: ‘upper carboniferous sediments’ (c2), the new unit ‘lower permian sediments, undivided’ (lp) which is only found at prinsesse ingeborg halvø, and the upper permian ‘midnatsfjeld and kim fjelde formations’ (mi). furthermore, a fault sliver of proterozoic basement of the ‘independence fjord group’ (if) has also been mapped (fig. 2b). updates based on recently published work on the kap rigsdagen beds following a recent publication identifying these beds as belonging to the upper unit of the early cretaceous ladegårdsåen formation (piasecki et al. in press), the outcrops found at kap rigsdagen (fig. 1) are also assigned to the new map unit lc. likewise, a small locality in southern kim fjelde in peary land is interpreted as lower cretaceous (håkansson et al. 1981). in previous versions of the 1:500 000 scale map (bengaard & henriksen 1986) these two areas were assigned to the very broad map unit jc (upper jurassic – lower cretaceous sediments) along with strata at kilen. this map unit is thus no longer necessary for the 1:500 000 scale map and has been emended. piasecki et al. (in press) further ascribed a new occurrence of the paleocene to possibly eocene thyra ø formation on top of the ladegårdsåen formation at kap rigsdagen. this occurrence has been shown as a small polygon with dashed outline on the updated map, since the lateral distribution of the formation at this locality is unknown (fig. 1, arrow). suggested procedure for updating geological maps at geus in order to make sure that the maps published by geus reflect the current state of geological knowledge, the digital geological maps should be routinely updated by including results from of multi-year regional field work such as the 2012, 2013 and 2016 expeditions to the wandel sea basin, and which was also done after field work in south-east greenland (kolb et al. 2016). in order to ensure the quality of the maps, it is important that only quality-tested and preferably peer-reviewed data 4242 are used for the updates and that there is a fixed procedure for how they are carried out and reported. a map editor or potentially an external reviewer, checks the map and legend for consistency and quality. before the update is carried out, the previous versions of the map are saved in the database and the changes to the map are recorded and documented in the metadata of the mapped elements. finally, the details of the changes are reported as a geus map sheet description or in a brief paper such as the present one. when an area corresponding to a map sheet (see holst et al. 2013, p. 59) has been sufficiently updated to justify printing of a new version, the updated map sheet can be printed along with a description. there is a significant backlog in updating geus’ series of map sheets with already published data. as exemplified above, systematic updating has not previously been the custom. conducting this is important to make sure that the official maps published by geus reflect the current state of geological knowledge. acknowledgements thanks are due to s. bernstein and g.k. pedersen for helpful suggestions to earlier versions of this manuscript. w. weng implemented the changes to the digital map. references alsen, p., mcroberts, c., svennevig, k., bojesen-koefoed, j.a., hovikoski, j. & piasecki, s. 2017: the isrand formation: a middle triassic daonella-bearing, black shale unit in kilen, north greenland (with a note on the triassic in amdrup land). newsletters on stratigraphy 50, 31–46. bengaard, h.-j. & henriksen, n. 1986: geological map of greenland, 1:500  000, peary land, sheet 8. copenhagen: geological survey of greenland. bojesen-koefoed, j.a., alsen, p. & christiansen, f.g. 2014: six years of petroleum geological activities in north-east greenland (2008–2013): projects and a view of the future. geological survey of denmark and greenland bulletin 31, 59–62. henriksen, n. 2003: caledonian orogen, east greenland 70°-82°n. geological map 1:1 000 000. a compilation of lithostructural data. copenhagen: geological survey of denmark and greenland henriksen, n., higgins, a.k., kalsbeek, f. & pulvertaft, t.c.r. 2009: greenland from archaean to quaternary. descriptive text to the 1995 geological map of greenland, 1:2  500  000. 2nd edition. geological survey of denmark and greenland bulletin 18, 126 pp. holst, j., glendal, e.w. & dawes, p.r. 2013: catalogue of survey publications on greenland. copenhagen: geological survey of denmark and greenland, 108 pp. hovikoski, j., pedersen, g.k., alsen, p., lauridsen, b.w., svennevig, k., nøhr-hansen, h., sheldon, e., dybkjær, k., bojesen-koefoed, j.a. & piasecki, s. 2018. in press: the jurassic–cretaceous lithostratigraphy of kilen, kronprins christian land, eastern north greenland. bulletin of the geological society of denmark 66, 61–114. håkansson, e., birkelund, t. & piasecki, s. 1981: jurassic-cretaceous boundary strata of the extreme arctic (peary land, north greenland). bulletin of the geological society of denmark 30, 11–42. jepsen, h.f. 2000: geological map of greenland 1:500  000, lambert land, sheet 9. copenhagen: geological survey of denmark and greenland. kokfelt, t.f., keulen, n., weng, w.l. & pedersen, m. 2013: geological map of greenland, 1 : 500  000. copenhagen: geological survey of denmark and greenland. kolb, j., steensgaard, b.m. & kokfelt, t.f. 2016: geology and mineral potential of south-east greenland. danmarks og grønlands geologiske undersøgelse rapport 2016/38, 1–157. pedersen, m., weng, w.l., keulen, n. & kokfelt, t.f. 2013: a new seamless digital 1:500  000 scale geological map of greenland. geological survey of denmark and greenland bulletin 28, 65–68. pedersen, s.a.s. & håkansson, e. 1999: kronprins christian land orogeny deformational styles of the end cretaceous transpressional mobile belt in eastern north greenland. polarforschung 69, 117–130. piasecki, s., nøhr-hansen, h. & dalhoff, f. 2018. in press: revised stratigraphy of kap rigsdagen beds, wandel sea basin, north greenland. newsletters on stratigraphy, 15 pp. http://dx.doi.org/10.1127/ nos/2018/0444 stemmerik, l., håkansson, e., madsen, l., nilsson, i., piasecki, s., pinard, s. & rasmussen, j.a. 1996: stratigraphy and depositional evolution of the upper palaeozoic sedimentary succession in eastern peary land, north greenland. grønlands geologiske undersøgelse bulletin 171, 45–71. stemmerik, l., larsen, b.d. & dalhoff, f. 2000: tectono-stratigraphic history of northern amdrup land, eastern north greenland: implications for the northernmost east greenland shelf. geology of greenland survey bulletin 187, 7–19. svennevig, k. 2018. in press: geological map of greenland, 1:100  000, kilen 81 ø.1 syd. copenhagen: geological survey of denmark and greenland. svennevig, k., guarnieri, p. & stemmerik, l. 2016: tectonic inversion in the wandel sea basin: a new structural model of kilen (eastern north greenland). tectonics 35, 2896–2917. svennevig, k., guarnieri, p. & stemmerik, l. 2017: 3d restoration of a cretaceous rift basin in kilen, eastern north greenland. norwegian journal of geology 97, 21–32. svennevig, k., alsen, p., guarnieri, p., hovikoski, j., lauridsen, b.w., pedersen, g.k., nøhr-hansen, h. & sheldon, e. 2018. in press: descriptive text to the geological map of greenland, 1:100 000, kilen 81 ø.1 syd. geological survey of denmark and greenland map series 8. author’s address geological survey of denmark and greenland (geus), østeer voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ksv@geus.dk. http://dx.doi.org/10.1127/nos/2018/0444 http://dx.doi.org/10.1127/nos/2018/0444 mailto:ksv@geus.dk geological survey of denmark and greenland bulletin 33, 2015, 69-72 69 automatic weather stations for basic and applied glaciological research michele citterio, dirk van as, andreas p. ahlstrøm, morten l. andersen, signe b. andersen, jason e. box, charalampos charalampidis, william t. colgan, robert s. fausto, søren nielsen and martin veicherts since the early 1980s, the geological survey of denmark and greenland (geus) glaciology group has developed automatic weather stations (awss) and operated them on the greenland ice sheet and on local glaciers to support glaciological research and monitoring projects (e.g. olesen & braithwaite 1989; ahlstrøm et al. 2008). geus has also operated awss in connection with consultancy services in relation to mining and hydropower pre-feasibility studies (colgan et al. 2015). over the years, the design of the aws has evolved, partly due to technological advances and partly due to lessons learned in the fi eld. at the same time, we have kept the initial goal in focus: long-term, year-round accurate recording of ice ablation, snow depth and the physical parameters that determine the energy budget of glacierised surfaces. geus has an extensive record operating awss in the harsh arctic environment of the diverse ablation areas of the greenland ice sheet, glaciers and ice caps (fig. 1). th e current geus-type aws (fig. 2) records meteorological, surface and sub-surface variables, including accumulation and ablation, as well as for example ice velocity. a large part of the data is transmitted by satellite near real-time to support ongoing applications, fi eld activities and the planning of maintenance visits. th e data have been essential for assessing the impact of climate change on land ice. th e data are also crucial for calibration and validation of satellite-based observations and climate models (van as et al. 2014). th e current version of the geus aws was developed in 2007 (sensors and tripod) and in 2008 (data logger, satellite data telemetry and power management) coinciding with the establishment of the programme for the monitoring of the greenland ice sheet (promice; ahlstrøm et al. 2008) and the glaciobasis programme monitoring an ice cap in a.p. olsen land in north-east greenland (apo; fig. 1). in con© 2015 geus. geological survey of denmark and greenland bulletin 33, 69–72. open access: www.geus.dk/publications/bull 22 50 stations on the ice sheet stations on local glaciers discontinued stations 250 km 2250 3000 2750 2500 1000 1500 1750 20 00 25 00 22 50 17 50 1500 1 2 5 0 2000 apo qas nuk kan mal thu upe kpc sco mit tas har nhf nas ime ami paq ste jdl fig. 1. sites monitored by the geological survey of denmark and greenland with automatic weather stations. the currently active sites on the greenland ice sheet (red dots) consist of transects with two or three stations at different elevations. qas: qassimiut (2000–). nuk: qamanaarsuup sermia, nuuk (1979–1989, 2008–). kan: kangerlussuaq (2008–). upe: upernavik (2008–). thu: tuto ramp, thule (2008–). kpc: kronprins christian land (1993–1994, 2008–). sco: violin gletscher near scoresby sund (2008–). tas: tasiilaq (2004–). currently active sites on local glaciers (blue dots), with one to three aws per site are apo: ice cap in a.p. olsen land (2008–). mal: malmbjerg (2008–). mit: mittivakkat glacier (1995–). sites where geus had automatic weather stations in the past (black dots), some of which in cooperation with the former greenland technical organisation: jdl: nordbogletscher, johan dahl land (1977–1983). ami amitsuloq ice cap (1981–1990). paq: pakitsoq (1984–1987). nas: narsap sermia (2003–2006). iso: isortuarsuup tasia (1984–1987). nhf: nioghalvfjerdsfjorden (1996–1997). har: hare gletscher (1994–1995). ste: steenstrup gletscher (2004–2008). ime: imersuaq (1999–2002). sts: storstrømmen (1989–1994). 7070 nection with consulting work, the fi rst aws with the new design (mal; fig. 1) was installed in 2008 for quadra mining ltd., vancouver, canada (now kghm international ltd., lubin, poland) near the malmbjerg molybdenum occurrence in stauning alper, central east greenland (citterio et al. 2009), followed in 2008 and 2009 by three more stations on the ice sheet in the kangerlussuaq region (kan, fig. 1). th ese stations are part of the greenland analogue project (gap; van as et al. 2012) for skb, the swedish nuclear fuel and waste management company (stockholm, sweden) and posiva oy (olkiluoto, eurajoki, finland). th e geus aws model in use now is a reliable tool that is adapted to the environmental and logistical conditions of polar regions. it has a proven record of more than 150 stationyears of deployment in greenland since its introduction in 2007–2008, and a success rate of c. 90% defi ned as the fraction of months with more than 80% valid air-temperature measurements over the total deployment time of the 25 stations in the fi eld. th e rest of this paper focuses on the technical aspects of the geus aws, and provides an overview of its design and capabilities. station requirements th e geus aws is the fundamental component of a monitoring network which can include numerous stations, a satellite data link, and a receiving database where telemetry data are decoded and validated before further analysis and dissemination. th e cost-eff ective aws delivers timely researchquality data year-round from glacier ablation areas in remote locations. th e aws must therefore require little maintenance, with a target of maximum one visit per year. power generation and battery capacity must be suffi cient to operate through the polar night. data quality must be assured by accurate measurement techniques including aspiration of radiation shields and tilt correction of (shortwave) radiometer measurements. th e mechanical construction of the station must keep the sensors at a constant height above the ice surface, which can ablate more than 9 m of ice per year in south greenland (van as et al. 2011), and the station must be able to survive burial in snow in the winter months. timeliness of data availability and the assessment of station health demand satellite data telemetry both in summer and winter. to our knowledge, no other commercially available aws satisfi es all these requirements. sensors, data logger and telemetry th e geus aws can be fi tted with any sensor, but the standard aws measures air temperature and humidity, wind speed and direction, atmospheric pressure, downward and refl ected solar shortwave radiation, downward and upward longwave radiation, subsurface (ice) temperatures, snow depth, ice ablation, gps position, as well as diagnostic parameters such as battery voltage and ventilator power consumption, and 2-axes station tilt necessary for correcting shortwave radiation measurements. table 1 lists sensor types, their measurement heights and uncertainty as specifi ed by the manufacturers. th e geus aws is designed to minimise measurement errors where possible, for instance by actively aspirating the radiation shield inside which air-temperature and humidity sensors are located. successful error detection and management increases with user experience and specialist knowledge. two of the sensors listed in table 1 are designed and manufactured in-house at geus for use on ice: the 8-levels thermistor string and the pressure transducer assembly (pta) that measure ice ablation (fausto et al. 2012). th e pta works by relating decrease of bottom hydraulic pressure (corrected for atmospheric pressure), measured inside an antifreeze-mixture fi lled hose drilled into the ice, to surface a cb d 10 9 8 7 11 25 6 14 3 12 fig. 2. standard sensor suite of the automatic weather station on arcturus gletscher at the malmbjerg molybdenum prospect. a: the main instrument boom. b: the tripod and the sites of the thermistor string and pressure-transducer assembly drilled into the ice. c: the sonic rangerstake frame drilled into the ice. d: the inside of the data logger enclosure. the numbers refer to the list in table 1. 71 lowering due to ice ablation. both the pta and the thermistor string can easily be interfaced to most data loggers. all analog and digital sensors are connected to a campbell cr1000 data logger housed in a watertight enclosure together with a campbell am16/32a analog multiplexer and supporting circuitry. th e logger is programmed to record in 10-minute cycles throughout the year. th e only exception is the gps, which is not needed at such a high rate and is activated less frequently in order to economise power. th e campbell cr1000 data logger is an established platform that is widely used in polar climates both in the arctic and in antarctica (lazzara et al. 2012). th e multiplexer is confi gured to support half-bridge measurement of thermistors from up to four 8-level thermistor strings (only one is normally used), in addition to 32 single-ended or 16 diff erential analog measurements (only six of each type are normally used), providing large fl exibility for customised sensor suites. th e main local data storage is a removable fl ashcard rated for operation over extended temperature ranges. for reference, a 256 mb card will log in excess of 7 years of 10-minute records. to provide redundancy of data storage, the internal logger memory is confi gured to store 1-hour average records and can hold in excess of one year of data as a backup for the fl ashcard. data can be retrieved during on-site maintenance by swapping the fl ashcard or downloading its content to a laptop. th e robustness of the system is illustrated in fig. 3 by an uninterrupted 2008–2015 time series from the apo_m station in north-east greenland, the fi rst aws built entirely according to the current geus design. aft er seven years in the fi eld, this station still employs the original electronics, telemetry, battery and tripod hardware although sensors have been periodically replaced according to a scheduled recalibration plan. satellite data telemetry can transmit up to 340 bytes per message through the iridium short burst of data (sbd) service. th e program running on the cr1000 data logger implements a full soft ware handshake with the transmitter to ensure that a satellite is in view and that data are correctly transferred from the logger to the transmitter and the iridium satellite. if no acknowledgement of successful transmission is received from the satellite, the data logger will retry the transmission once, or queue the unsent message for delivery at a later time, depending on iridium service availability. th is mode of operation ensures a low rate of message loss and relatively low power consumption by avoiding unnecessary transmission attempts. th e stations binary-encode data before transmission, reducing data transmission costs by about 2/3. further transmission costs can be saved by transmitting the less transient variables at longer intervals. power th e long polar night and low temperatures exert a strong infl uence on the aws design. th e aws operates on solar power and rechargeable sealed lead-acid batteries for a nomi fig. 3. an example of 10-minute observations of air temperatures from 2008 to 2015 from the apo_m station on an ice cap in a.p. olsen land in north-east greenland. table 1. current sensors used on a standard geus automatic weather station parameter and sensor height manufacturer, type and sensor accuracy the numbers refer to fig. 2 where the positions of the sensors on an automatic weather station are shown.the heights of the sensors above the surface are indicative. the accurate heights are measured before and after every maintenance visit. 1. air temperature and relative rotronic mp102h with pt100 and hc2-s3 probe (±0.1 k, ± 0.8% rh, at humidity, 2.7 m 23°c ± 5 k), housed in a rs12t aspirated shield 2. radiation (downward and reflected solar shortwave, kipp & zonen cnr1 (uncertainty in daily totals < 10%) or cnr4 downward and upward longwave; 2.9 m) (uncertainty in daily totals < 5% shortwave, < 10% longwave) 3. wind speed and direction, 3.1 m r.m. young 05103-5 (± 0.3 ms–1 or 1%, ± 3°, non-riming conditions) 4. snow depth campbell scientific sr50 or sr50a (± 1 cm or 0.4%) 5. 2-axes radiometer tilt hl planar ns-25/e2 in geus assembly (± 0.2°) 6. iridium satellite antenna iridium at1621-142 quad-helix 7. iridium sbd modem u-blox neo module 7. iridium sbd modem nal research sbd-9601 or sbd-9602 8. atmospheric pressure setra model 278 (± 2.5 hpa, at −40 to 60°c) 9. ice ablation and snow depth campbell scientific sr50 or sr50a (± 1 cm or 0.4%) 10. subsurface temperature profile geus thermistor string with 8 rs components thermistors 151-243 (± 0.2°c, at 0°c) 11. ice ablation geus pta with ørum & jensen elektronik nt1400 or nt1700 (± 2.5 cm) 12. gps antenna trimble p/n 56237-40 active ceramic patch 29 july 2010 29 july 2012 29 july 2014 10 0 −10 −20 −30te m p er au re ( °c ) 29 july 2008 7272 nal total of 112 ah at 12 v. th e power system is composed of the main unregulated 12 v power rail permanently supplying the data logger directly from the batteries, two 12 v unregulated rails controlled by soft ware through two external solid-state switches, and one switched 5 v regulated rail under direct data logger control. th is arrangement allows independent powering up of the radiation shield aspirator fan, the gps and satellite transmitter, the 5 v loads of the sonic rangers, temperature and humidity sensors, tilt meter and multiplexer. a single 10 w solar panel is wired to the main 12 v rail through a power schottky diode to prevent that the solar panel drains the battery during the winter months, and to eliminate the need for a charge regulator, which occasionally failed in previous aws designs. a soft ware-controlled low-power mode is activated when battery voltage under load falls below a confi gurable threshold (set to 11.5 v), which is never reached in normal circumstances. in low-power mode, operation continues almost as normal, but the most power-demanding functions (aspiration fan, satellite telemetry and gps) are deactivated. th e low-power mode is exited once solar charging brings the battery suffi ciently above the voltage threshold. th e soft ware can be confi gured for polar day and night operation, for instance to reduce data transmission rates during winter. th e typical monthly power requirements of a geus aws as confi gured for promice is 17 ah in summer, 11 ah in winter and 1.3 ah in low-power mode. tripod and sonic ranger frame th e tripod is constructed from 1˝ and 1.5˝ aluminium tubes with steel wires connecting legs and mast in a tetrahedral structure for a stable free-fl oating tripod. most of the sensors are fi xed to a horizontal boom at c. 2.9 m above surface. th e battery box, which weighs c. 50 kg, is suspended under the mast to improve station stability by increasing the aws mass and lowering the centre of gravity. th e tripod can be folded and transported in a small helicopter. during maintenance visits, which normally take 3–4 hours and include replacements of sensors due for recalibration, re-drilling of sensors and occasional repairs, the tripod can be easily tilted so that it does not have to be disassembled. th e sonic ranger frame is also built from 1˝ aluminium tubes, and its three vertical legs are drilled into the ice a few metres away from the aws tripod. concluding remarks th e geus aws has been developed, produced and deployed operationally by geus, and supplied to partners within denmark and abroad. it is a proven solution for a wide range of basic and applied glaciological research in arctic and alpine settings and is available through research collaborations or commercial sale. th e standard design can accommodate signifi cant expansion of the sensor suite. th e geus aws is readily available and supported as a stand-alone or as a component of wider services including fi eld deployment, maintenance, training and data management and analysis. acknowledgements we thank ole b. olesen, carl e. bøggild, wim boot, peer jørgensen, and other people who were instrumental in the technical development of the geus aws over the years. the promice and glaciobasis monitoring programmes are funded by the danish energy agency (ens). references ahlstrøm, a.p. and the promice project team 2008: a new programme for monitoring the mass loss of the greenland ice sheet. geological survey of denmark and greenland bulletin 15, 61–64. citterio, m., mottram, r., larsen, s.h. & ahlstrøm, a.p. 2009: glaciological investigations at the malmbjerg mining prospect, central east greenland. geological survey of denmark and greenland bulletin 17, 73–76. colgan, w., thomsen, h.h. & citterio, m. 2015: unique applied glaciology challenges of proglacial mining. geological survey of denmark and greenland bulletin 33, 61–64. fausto, r.s., van as, d., ahlstrøm, a.p. & citterio, m. 2012: assessing the accuracy of greenland ice sheet ice ablation measurements by pressure transducers. journal of glaciology 58, 1144–1150. lazzara, m.a., weidner, g.a., keller, l.m., thom, j.e. & cassano, j.j. 2012: antarctic automatic weather station program: 30 years of polar observation. bulletin of the american meteorological society 93, 1519–1537. olesen, o.b. & braithwaite, r.j. 1989: field stations for glacier-climate research, west greenland. in: oerlemans, j. (ed.): glacier fluctuations and climatic change, 207–218. dordrecht: kluwer academic publishers. van as, d., fausto, r.s. & the promice project team 2011: programme for monitoring of the greenland ice sheet (promice): first temperature and ablation records. geological survey of denmark and greenland bulletin 23, 73–76. van as, d., hubbard, a.l., hasholt, b., mikkelsen, a.b., van den broeke, m.r. & fausto, r.s. 2012: large surface meltwater discharge from the kangerlussuaq sector of the greenland ice sheet during the recordwarm year 2010 explained by detailed energy balance observations. the cryosphere 6, 199–209. van as, d. et al. 2014: increasing meltwater discharge from the nuuk region of the greenland ice sheet and implications for mass balance (1960–2012). journal of glaciology 60, 314–322 authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: mcit@geus.dk research article rosa et al. 2020: geus bulletin 44. 4626. https://doi.org/10.34194/geusb.v44.4626 1 of 8 petrography, geochemistry and magnetic susceptibility of the isortoq fe-ti-v deposit, isortoq giant dykes, south greenland diogo rosa*, alessandro sandrin, troels f.d. nielsen, høgni vesturklett department of petrology and economic geology, geological survey of denmark and greenland (geus), copenhagen, denmark abstract the isortoq giant dykes in the proterozoic gardar province, south greenland, include the isortoq south giant dyke and the isortoq north giant dyke. the fine-grained fe-ti-v deposit hosted by the isortoq south giant dyke, referred to as the isortoq fe-ti-v deposit, is considered a good test site for the use of magnetic susceptibility for the mapping of ore grades. here, we test this and show that the fe, ti and v distribution is controlled by titanomagnetite disseminated throughout fine-grained troctolite. the deposit displays a clear correlation between magnetic susceptibility and fe, ti and v grades in bulk samples of consecutive 2 m sections from 11 drill cores, totalling 2671 m in length. we observe that fe, ti and v are almost entirely hosted in titanomagnetite, which controls the magnetic susceptibility. field measurements of the magnetic susceptibility can thus be considered as a reliable exploration tool for this type of mineralisation. we further consider the origins of the deposit by reconnaissance petrography, mineral and bulk rock chemistry of the large mass of aphanitic fe-rich troctolite in the isortoq south giant dyke. we suggest that the deposit may represent the base of a basanitic to trachybasaltic magma chamber, in which fe-rich immiscible melts accumulated, crystallised and fractionated. the processes suggested here may apply to other giant dykes and intrusions of the gardar province. 1 introduction the isortoq iron-titanium-vanadium (fe-ti-v) deposit is located within the isortoq giant dykes system (60.96°n, 47.43°w; fig. 1) in the proterozoic gardar province, south greenland (upton 2013). the deposit is estimated to hold a resource of 70 million tonnes with an average of 29.6% fe, 10.9% tio2, and 0.144% v2o5, at a cut-off of 15% fe (turner & nicholls 2013). these metals are used by the steel, pigment and battery industries; and thus, the isortoq fe-ti-v deposit is of interest to exploration companies. and yet its petrophysical parameters, needed for the inversion of magnetic surveys, are poorly understood, and the origin of the deposit remains unclear. *correspondence: dro@geus.dk received: 10 mar 2020 accepted: 06 may 2020 published: 21 aug 2020 keywords: isortoq fe-ti-v deposit, gardar province, petrophysics, magma chamber, troctolite abbreviations: ms: magnetic susceptibility ogdc: older giant dyke complex ygdc: younger giant dyke complex emp: electron microprobe tas: total alkalis silica geus bulletin is an open access, peerreviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: mette olivarius (geus, denmark) reviewed by: tod waight (university of copenhagen, denmark) and claire nichols (massachusetts institute of technology, usa) funding: see page 8 competing interests: none declared additional files: see page 8 https://doi.org/10.34194/geusb.v44.4626 mailto:dro@geus.dk rosa et al. 2020: geus bulletin 44. 4626. https://doi.org/10.34194/geusb.v44.4626 2 of 8 www.geusbul let in.org the ore is a very fine-grained troctolite with high concentrations of disseminated feti-oxides in a dykelike intrusive body that is up to 200 m wide, 230 m deep and >6 km long. the troctolite has no clear cumulus textures and is considered unusual because the grains are too small to sink in the magma. therefore, the high fe in the deposit cannot be because of the accumulation of titanomagnetite. if the high content of fe-ti-v is not because of the accumulation of titanomagnetite, then high fe-ti-v must be because of the accumulation of fe-rich melt from which titanomagnetite crystallised. our investigation, therefore, has two aims: (1) to test the use of magnetic susceptibility (ms) for mapping of the deposit (sandrin & elming 2006, 2007; sandrin et al. 2009) and (2) to provide a hypothesis for the deposit’s origin. the study is based on the data obtained from the diamond core during two drilling programmes of the isortoq fe-ti-v deposit (fig. 1), namely four samples for petrographic studies from core sl306 (drilled in 2005 by hunter minerals; ferguson 2010) and 11 cores (dh1–11, drilled in 2012 by west melville metals inc.; ferguson 2013; fig. 1), and ms data produced for this study. a strong correlation between core lithology, chemistry and ms would validate ms as a reliable, fast and lowcost tool for field mapping of fe, ti and v distribution elsewhere in the isortoq giant dykes system and in the gardar province. it would also help to constrain areas that could be studied for grain size, morphology and mineralogy using the micro-beam instrumentation and provide important petrophysical parameters for inversion of magnetic surveys. in this study, we have systematically measured ms along a total of 2671 m in cores dh1–11 and correlated to the bulk compositions of a total of 1069 samples from the same drill cores (ferguson 2013). in addition, we include information on petrography, bulk rock and mineral chemistry from all 12 cores and present a conceptual model for the processes in the giant dyke that led to the accumulation of magnetite and, consequently, fe, ti and v enrichment. these insights constitute key parameters for exploration, evaluation and genesis of this deposit and potentially elsewhere in the gardar province. 2 geological setting the isortoq giant dykes consist of (1) isortoq south giant dyke and (2) isortoq north giant dyke. these dykes were emplaced in older gneiss and granite of the julianehåb batholith of the ketilidian mobile belt of south greenland (pulvertaft 2008). the giant dykes are part of the middle proterozoic gardar rift and magmatic province comprising lavas, dyke swarms, intrusive complexes and the older giant dyke complex (ogdc) and younger giant dyke complex (ygdc) of the tuttutooq area (upton 2013). the isortoq south giant dyke extends over a distance of >6 km, while the isortoq north giant dyke can be traced for c. 8.5 km to the edge of the greenland ice sheet. the 12 cores used in this study were drilled in the main body of the isortoq south giant dyke, sw of snoopy lake (fig. 1). the large tabular intrusive bodies are referred to as giant dykes. however, all but one of the 12 cores drilled in the isortoq south giant dyke exit into granitic host rock and show that isortoq south is a tabular intrusive body with a floor contact at 100–200 m depth, fig. 1 geological map of isortoq south giant dyke, south greenland. yellow circles: location of drill cores. yellow lines: projection to surface (for non-vertical drill holes) for the 12 studied drill cores. black dashed lines: dyke extension. solid red line: extent of mineral resource established by turner & nicholls (2013). inset map: includes the divisions of the ketilidian orogen of chadwick & garde (1996). red circle: location of the study area. green circle: tuttutooq. https://doi.org/10.34194/geusb.v44.4626 http://www.geusbulletin.org rosa et al. 2020: geus bulletin 44. 4626. https://doi.org/10.34194/geusb.v44.4626 3 of 8 www.geusbul let in.org dependent on collar elevation of the drill hole. the contact at the wall between the granitic host rocks and the isortoq south giant dyke is composed of chilled basalt or fine-grained troctolite, whereas the contact at the floor is characterised by mingling with rheomorphic melts derived from the host rocks (core photos are provided in supplementary file si 1). 3 samples and methods core sl306 was drilled by hunter minerals in 2005. in 2006, four samples were provided to the geological survey of denmark and greenland (geus) for electron microprobe (emp) analyses (see supplementary file s2). the main silicate, feti-oxides and rare sulphides were analysed for major and minor elements using a jeol jxa-8200 superprobe emp at the department of geosciences and natural resource management, university of copenhagen, geocenter, denmark. the following elements were analysed (sio2, tio2, al2o3, cr2o3, v2o3, feo, mno, nio, mgo, cao, na2o, k2o in silicates and fetioxides, and mn, s, cu, fe, ni and co in sulphides). counting times were 20 s on peaks and 10 s on backgrounds, following standard procedures. the acceleration voltage was 15 kv, and the beam current was 15 na. beam diameters were 5 μm. a combination of pure metal and mineral standards was used. in 2011, west melville metals inc. drilled 11 cores (dh1–11) for systematic assays, petrographic and geochemical studies. grain size and whole-rock compositions for a total of 1069 samples were reported by ferguson (2013). the data were provided for use here by resource500 fevti – the present holder of the exploration licence. the core was split along its length and divided into 2 m sections. all material within each 2 m section was pooled for analysis. as such, each sample represents the average composition over each consecutive 2 m section of core, with no hiatuses in the sampling. major elements (sio2, al2o3, fe2o3(t), mno, mgo, cao, na2o, k2o, tio2, p2o5, cr2o3 and v2o5) were analysed by the fused puck xrf analysis. methods are described in ferguson (2013). as part of the original analytical work, trace elements (ag, cd, cu, mn, mo, ni, pb, zn, be, bi, co, sr and y) were analysed by four acid (near total) digestion followed by icp-ms analysis at actlabs, canada. these data are available in ferguson (2013), but are not discussed here. fig. 2 textural and mineralogical aspect of the troctolite. a: cross section of drill core from dh3 (sample 12 dtq003; depth c. 53 m). photo courtesy of resource 500 fevti ltd). note the fine-grained and featureless texture of the troctolite. b: photomicrograph (transmitted light) of sample 12 dtq003 in dh3 (depth c. 53 m) showing olivine (ol), chlorite (chl), biotite (bt), plagioclase (plag), hornblende (hbl), titanomagnetite (ti mag) and richterite amphibole (rich). photo courtesy of resource 500 fevti ltd. note the amphiboles partially replace clinopyroxene crystals. c: backscatter image of sample from core sl 306 (depth c. 62 m), showing magnetite (mag, light grey), ilmenite (ilm, grey) and other rock minerals such as pyroxene, olivine and plagioclase as dark to black. note the fine-grained nature of the ore minerals. small sulphide grains (white) are attached to titanomagnetite grains. https://doi.org/10.34194/geusb.v44.4626 http://www.geusbulletin.org rosa et al. 2020: geus bulletin 44. 4626. https://doi.org/10.34194/geusb.v44.4626 4 of 8 www.geusbul let in.org we measured ms on cores dh1–11, using a kt-10 ms meter, jointly designed by terraplus inc (canada) and georadis s.r.o. (czech republic). this handheld instrument uses an oscillator with an inductive coil to measure the ms and can be easily used in the field as well as on rock samples in the laboratory. drill cores dh1–11 all have a diameter of 3.65 cm. each ms reading represents a volume of core of c. 50 cm3. 4 results 4.1 petrography and mineral compositions (emp) petrographic characteristics for the isortoq south troctolite are illustrated in fig. 2. the grain size of the troctolite as recorded in cores dh1–11 is <500 μm with an average of c. 200 μm. the troctolite is quite homogenous and aphanitic in texture, though we observed 1 to 10 m thick cyclic units of troctolite with felsic tops near the base of the body (see supplementary file si 1). mineral compositions for core sl306 are presented in supplementary file si 2. the liquidus paragenesis includes olivine (fo67–62), clinopyroxene (magnesium number (mg#) 70–75; where mg# = mg/(mg + fe2+) × 100) and titanomagnetite. the liquidus phases are enclosed in plagioclase (andesine: from an51, or3 to an33, or7), biotite and very minor accessory pyrrhotite and pentlandite (supplementary file si 2). the troctolite was subjected to autometasomatism causing replacement of euhedral olivine by chlorite and magnetite. hornblende and richterite replace parts of clinopyroxene crystals (fig. 2b). in sl 306, titanomagnetite with limited exsolution of ilmenite has 18–22 wt.% tio2 and 0.2–0.5 wt.% v2o5, (fig.  2c). magnetite in titanomagnetite grains with significant volumes of exolved ilmenite has 2–8 wt.% tio2 and 0.6–0.8 wt.% v2o5 (data in supplementary file si 2). the elevated concentrations of v2o5 in the latter grains reflect the partitioning of v to the reduced volume of magnetite that was left after exsolution of ilmenite. 4.2 bulk rock compositions (xrf) major element data for all samples from drill cores dh1– 11 are presented in fig. 3 and table 1. sio2 is strongly negatively correlated with tio2 and feo* (*all  fe is fig. 3 average bulk rock analyses for consecutive 2 m intervals of drill core in cores dh1–11. a: sio2 vs feo*. b: sio2 vs tio2. c: sio2 vs mgo. d: sio2 vs na2o. a strong correlation exists between sio2 and tio2, feo*, mgo and na2o. note the control by composition at c. 16 wt.% sio2 and 50 wt.% feo*. average compositions for composition groups 1, 2 and 3 (see table 1) are shown. red dot: group 1, basanitic feeder, isortoq south dyke. orange dot: group 2, trachybasalt. yellow dot: group 3, basalts. dark blue square: average composition of basanitic chilled margins of older giant dykes. light blue square: average composition of trachybasaltic chilled margins of younger giant dykes. https://doi.org/10.34194/geusb.v44.4626 http://www.geusbulletin.org rosa et al. 2020: geus bulletin 44. 4626. https://doi.org/10.34194/geusb.v44.4626 5 of 8 www.geusbul let in.org reported as feo) and positively correlated with na2o (fig. 3). correlations are strongest below 42 wt.% sio2. we identify three compositional clusters (groups 1, 2 and 3; table 1 and fig. 3) between 42 and 50 wt.% sio2. sio2 exceeds 50 wt.% across the contact and into the host granites and gneisses of the isortoq south giant dyke. all drill cores, except for dh10, penetrated the floor into a felsic basement. drill core dh10 continued after a compositional break at c. 240 m from a troctolite with >30 wt.% feo* into basanite (total alkalis silica [tas] diagram, sensu le maitre et al. 2002) with <20 wt.% feo* and reached a depth of 344 m. the basanitic composition (group 1) with <20 wt.% feo* from below the general elevation of the floor of the troctolite body is interpreted as the possible feeder for the magma in the isortoq south giant dyke. the average composition of this feeder is comparable to the chilled margins of the ogdc at tuttutooq (upton 2013; table 1), although it is somewhat more fe-rich, which may be because of minor accumulation of titanomagnetite. group 2 and 3 decrease in alkalinity to compositions similar to those of the ygdc of tuttutooq (upton 2013; table 1) and have broadly comparable compositions of basaltic dykes in the gardar province (e.g. bartels et al. 2015). 4.3 magnetic susceptibility ms measured in dh 11 is shown in fig. 4. this drill core is presented here as the representative of all 11 drill cores (dh1–dh11) from the isortoq south giant dyke (diogo rosa, unpublished data 2019). also shown are the lithological variations between troctolite and granitic host rocks, as reported by west melville metals inc. table 1 average compositions for compositional groups 1, 2 and 3 of the isortoq south giant dyke and the giant dykes of tuttutooq   group 1 chill group 2 chill group 3   averagea ogdcb averagec ygdcd averagee isortoq tuttutooq isortoq tuttutooq isortoq   (%) (%) (%) (%) (%) sio2 43.58 44.63 46.48 46.69 48.2 tio2 4.79 4.52 3.27 2.67 2.19 al2o3 13.48 16.07 15.48 16.96 16.51 feo* 18.48 14.39 15.28 13.62 13.38 mno 0.25 0.2 0.21 0.19 0.19 mgo 5.89 4.89 5.62 6.02 6.07 cao 6.49 7.91 7.69 7.9 8.42 na2o 3.79 3.54 3.84 3.6 3.31 k2o 1.89 1.86 1.42 1.47 1.23 p2o5 1.36 2 0.72 0.87 0.49 sum 100 100 100 100 100 all fe as feo and classified according to le maitre et al. 2002 abasanitic feeder (drill core dh10, 242.00 to 343.88 m; ferguson 2013) bbasanitic chill of the older giant dyke complex (ogdc; upton 2013) ctrachybasalt average (45.85 to 47.20 wt% sio2; fig. 5 in this paper; data in ferguson 2013) dtrachybasaltic chill of younger giant dyke complex (upton 2013) etrachybasalt average (47.51 to 48.51 wt% sio2; see supplementary file si 2 for this paper; data in ferguson 2013) fig. 4 magnetic susceptibility (blue dots) and fe2o3, tio2 and v2o5 grades (orange dots) vs. drill depth in core dh11. vertical green bar: troctolite. vertical pink bar: host granitic rock. https://doi.org/10.34194/geusb.v44.4626 http://www.geusbulletin.org rosa et al. 2020: geus bulletin 44. 4626. https://doi.org/10.34194/geusb.v44.4626 6 of 8 www.geusbul let in.org (ferguson 2013), and the 2 m bulk rock averages for fe2o3, tio2 and v2o5 (ferguson 2013). ms readings were taken at 1 m intervals, whereas the geochemical data are average compositions for consecutive 2 m intervals along the core. thus, ms data were averaged (every two to three ms readings) to match the sampling of the geochemistry data. we correlate ms with concentrations of each of the three oxides (fe2o3, tio2 and v2o5), as well as between the oxides themselves. correlations between ms and the three oxides are very high (r = 0.96–0.97), as are the correlations between the oxides (r = 0.98–0.99; fig. 5). 5 discussion 5.1 the nature of the occurrence the isortoq fe-ti-v deposit was previously suggested to represent a down-faulted block of a lopolithic intrusion (ferguson 2010), in which the mineralisation formed by classic accumulation of titanomagnetite. however, the chilled margins of the deposit, the mingling with melted country rock, and the cyclic development in the floor all suggest the deposit is an intrusive body formed in a magma chamber bound by faults of the gardar graben system. the presence of a floor in 10 out of 11 drill cores and the occurrence of a possible feeder with a composition comparable to that of the ogdc at tuttutooq (table 1) in the deepest parts of dh 10 supports that the occurrence represents the lower most part of an elongated dyke-like magma chamber. 5.2 ms as exploration tool the ms investigation was initiated to test the use of ms as an exploration tool for the isortoq-type deposit. the ms of a rock is controlled by the type and modal proportion of the rock-forming minerals. often this control is exerted by the presence of relatively small quantities of ferromagnetic minerals, for example, magnetite, pyrrhotite (hrouda et al. 2009; sandrin et al. 2009). the logs in fig. 4 show that the troctolite intervals with the highest fe2o3 have one to two orders of magnitude higher ms values (0.3–0.4 international system of units (si) than those measured for granitic specimens. ms is strongly correlated with fe2o3 (r = 0.9578), tio2 (r = 0.9550) and v2o5 (r = 0.9719; fig. 5) in the central part of the deposit and confirms that ms is a valid exploration tool (caira 2012; ferguson 2013; turner & nicholls 2013). for most of the drill cores (diogo rosa, unpublished data 2019), inflexion points at ms values of approximately 0.05 si define the transition from granite to troctolite and mark a change in regression slopes (fig. 5). 5.3 towards a petrogenetic model for the isortoq fe-ti-v deposit the bulk rock analyses available in exploration reports give average compositions for consecutive 2 m intervals of drill core. the analysed bulk samples may, therefore, include granophyric veins or overlap with the host rock. a detailed account of the geochemistry and genesis of the deposit is, therefore, not possible without re-sampling. throughout the main body, the bulk rock compositions are controlled by the proportions of basanitic to trachybasaltic melt components, diluted by a fe-rich component with c. 16 wt.% sio2, 15 wt.% tio2, 50 wt.% feo* and 10 wt.% mgo (fig. 3). scatter in sio2 versus mgo (fig. 3) probably reflects minor variations in the titanomagnetite to olivine plus clinopyroxene ratios in the analysed samples and minor mineral settling within the troctolite. the basanitic composition (group 1 in table 1) and the fe-rich, thick cyclic units with felsic tops at the floor, suggest internal fractionation in semi-isolated fig. 5 correlations between magnetic susceptibility (ms) and bulk rock compositions in core dh11. a: ms vs. v2o5, b: ms vs. tio2. c: ms vs. fe2o3. https://doi.org/10.34194/geusb.v44.4626 http://www.geusbulletin.org rosa et al. 2020: geus bulletin 44. 4626. https://doi.org/10.34194/geusb.v44.4626 7 of 8 www.geusbul let in.org stratigraphic intervals. the combination of the finegrained texture and the absence of clear cumulative textures caused by gravitational sorting, and the formation of layers of titanomagnetite and olivine suggest that the fe-rich rocks of the isortoq south have crystallised in situ from dense and ponded fe-rich melt. upton (2013), building on bridgwater (1967), developed models for the magma chambers of the gardar province. upton proposed a model that includes crystallisation, fractionation and floatation of plagioclase under the roof of the emplaced magma, as well as ponding of dense crystal mushes in giant dykes. in addition to these processes, we suggest that magma in the top of the magma chamber evolved to trachyandesitic compositions and reached the twoliquid field between coexisting fe-rich and si-rich melts ( charlier & grove 2012). as such, we suggest a preliminary conceptual model for the origin of the isortoq fe-ti-v deposit (fig. 6). the most evolved trachyandesitic dyke compositions of the gardar region straddle the boundary to the immiscibility field (supplementary file si 3). fe-rich and dense immiscible melts segregated in a continuum and sank along the walls to the floor. in the isortoq south giant dyke we only the see the base of the magma chamber in which the fe-rich immiscible melts accumulated, crystallised and fractionated (fig. 6). this preliminary model is reminiscent of processes recently suggested for the skaergaard intrusion (nielsen et al. 2020). 6 conclusions strong correlations between ms and fe, ti and v concentrations confirm that ms is a valid tool for field mapping of the fe, ti and v distribution in the isortoq giant dykes system. further, the fe-ti-v mineralisation in the dyke is seen as the result of ponding and in situ crystallisation of immiscible fe-rich silicate melt. the immiscible melt formed when basanitic to trachybasaltic melts common to the gardar province fractionated and reached the two-liquid field between fe-rich and si-rich melts. acknowledgments the authors would like to thank resource500 fevti for allowing the publication of some of the results included in a confidential report. they also like to acknowledge nima azad for assisting with the ms readings. finally, this article benefited from the insights and discussions with stefan bernstein, thomas find kokfelt and jakob keiding (geus). additional information funding statement part of this work was financed by resource500 fevti as part of a study of the magnetic susceptibility of the diamond core drilled by west melville metals inc. author contributions dr: study coordination, exploration of the whole rock geochemistry data and its relationship with the ms, assessment of suitability of using ms for establishing grades. as: interpretation of the ms results. tfdn: carried out emp analysis and petrographic study, developed the petrogenetic model proposal. hv: performed ms readings on drill core. additional files three supplementary files are available alongside the article at https:// doi.org/10.34194/geusb.v44.4626 references bartels, a., nielsen, t.f.d., lee, s.r. & upton, b. 2015: petrological and geochemical characteristics of mesoproterozoic dyke swarms in the gardar province, south greenland: evidence for a major sub-continental lithospheric mantle component in the generation of the magmas. mineralogical magazine 79(4), 909–939. https://doi.org/10.1180/ minmag.2015.079.4.04 bridgwater, d. 1967: feldspathic inclusions in the gardar igneous rocks of south greenland and their relevance to the formation of major anorthosites in the canadian shield. canadian journal of earth sciences 4, 995–1014. https://doi.org/10.1139/e67-068 caira, n. 2012: ni43-101 technical report on the iron-titanium-vanadium potential of the isortoq property, south greenland, a territory of denmark. prepared by argonaut gold odysseys inc. on behalf of fig. 6 conceptual model for the formation of the isortoq south giant dyke. basanitic melt was emplaced in a magma chamber controlled by faults of the gardar graben and a floor of granitic host rocks. crystallisation and fractionation under the roof of the magma chamber resulted in floatation of feldspar crystals and evolution of the mush magma to the two-liquid field between feand si-rich silicate melts. dense fe-rich melts were continuously segregated and sank to the floor of the magma chamber where they crystallised and fractionated. https://doi.org/10.34194/geusb.v44.4626 http://www.geusbulletin.org https://doi.org/10.34194/geusb.v44.4626 https://doi.org/10.34194/geusb.v44.4626 https://doi.org/10.1180/minmag.2015.079.4.04 https://doi.org/10.1180/minmag.2015.079.4.04 https://doi.org/10.1139/e67-068 rosa et al. 2020: geus bulletin 44. 4626. https://doi.org/10.34194/geusb.v44.4626 8 of 8 www.geusbul let in.org west melville metals inc., 88 pp. http://www.aimva.com.au/literatureretrieve.aspx?id=185310 chadwick, b. & garde, a.a. 1996: palaeoproterozoic oblique plate convergence in south greenland: a re-appraisal of the ketilidian orogen. in: brewer, t.s. 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rift, south greenland. geological survey of denmark and greenland bulletin 29, 124 pp. https://doi.org/10.34194/geusb. v29.4692 https://doi.org/10.34194/geusb.v44.4626 http://www.geusbulletin.org http://www.aimva.com.au/literatureretrieve.aspx?id=185310 http://www.aimva.com.au/literatureretrieve.aspx?id=185310 https://doi.org/10.1144/gsl.sp.1996.112.01.10 https://doi.org/10.1007/s00410-012-0723-y https://doi.org/10.1007/s00410-012-0723-y https://data.geus.dk/gg_detail/?cat=rap&id=86591 https://data.geus.dk/gg_detail/?cat=rap&id=86591 https://data.geus.dk/gg_detail/?cat=rap&id=89801 https://data.geus.dk/gg_detail/?cat=rap&id=89801 https://www.geomatrix.co.uk/cms/resources/downloads/the-use-of-magnetic-susceptibility-of-rocks-in-geological-exploration-v1-1.pdf https://www.geomatrix.co.uk/cms/resources/downloads/the-use-of-magnetic-susceptibility-of-rocks-in-geological-exploration-v1-1.pdf https://www.geomatrix.co.uk/cms/resources/downloads/the-use-of-magnetic-susceptibility-of-rocks-in-geological-exploration-v1-1.pdf https://doi.org/10.1017/cbo9780511535581 https://doi.org/10.1017/cbo9780511535581 https://doi.org/10.1093/petrology/egz057 https://doi.org/10.1093/petrology/egz057 https://doi.org/10.1016/j.gexplo.2009.07.002 https://doi.org/10.1016/j.gexplo.2009.07.002 https://doi.org/10.1016/j.oregeorev.2005.06.001 https://doi.org/10.1016/j.oregeorev.2006.02.002 https://doi.org/10.1016/j.oregeorev.2006.02.002 https://data.geus.dk/gg_detail/?cat=rap&id=93978 https://data.geus.dk/gg_detail/?cat=rap&id=93978 https://doi.org/10.34194/geusb.v29.4692 https://doi.org/10.34194/geusb.v29.4692 geological survey of denmark and greenland bulletin 31, 2014, 95-98 95 to what extent is denmark vulnerable to mineral supply shortage? per kalvig, rune j. clausen, niels fold and karen hanghøj mineral resources are building blocks of modern society and essential for progress and prosperity. mankind has always depended on access to mineral raw materials, which have been a key factor for wealth, culture and development. modern societies are characterised by a rapidly increasing demand for specialised mineral raw materials, determined by their stage of technological development, the number of consumers, and their standard of living. generally, the availability of mineral raw materials has not, until recently, been considered an issue by the average consumer or by companies in the downstream end of the value chains, and mineral resources have not been part of the political agenda. in this context china’s control over rare-earth elements (ree) has been an eye opener to both industry and politicians worldwide, and has subsequently led to discussions about the possible exhaustion of finite resources and potential threats to the availability of raw materials caused by geopolitical tension and market restrictions. the increased concern has lead to several attempts to assess the risk of supply shortage which are however still at a rather qualitative stage. inadequate knowledge about the current and future demand for mineral raw materials prevents political and industrial decision-makers from taking the necessary actions to predict and mitigate the national and industrial vulnerability to supply shortage. thus, most modern societies, including denmark, are vulnerable to mineral raw materials scarcity, but unaware of where and how it may appear, and how to prevent and address the problem. scarcity issues scarcity issues have been discussed since thomas malthus in 1798 initially predicted problems of food shortage due to increasing population and later also in relation to mineral resources. a number of organisations and individuals (e.g. club of rome, gro harlem bruntland) have taken the lead in these discussions and emphasised that natural resources are finite and limited and that the global economy is growing disproportionately. at the summit meeting in rio de janeiro in 1992, all nations were encouraraged to adopt the so-called bruntland principles to ensure sufficient resources for future generations. the term sustainability was introduced to the mining industry. however, no clear effects can be identified neither in the policies nor in the overall mineral consumption, and global and national concern on how to secure raw material supply is increasing. terms such as critical minerals were introduced, reflecting the risk of scarcity of some raw materials. the us national research council quantitatively addressed scarcity issues related to minerals in 2008 (national research council 2008), and since then a substantial number of reports have focused on the topic (e.g. rosenou-tornow et al. 2009; european commission 2010; undp 2010; graedel et al. 2012). the decoupling of wealth and mineral resource consumption remains to be seen. why are minerals important? mineral-based materials are present everywhere in our daily life – in houses, cars, computers, cooking utensils, paint, tiles, paper, plastic, batteries, wind turbines, roads, pipes etc. for each and all of these ‘end products’ the choice of raw materials – and thus the minerals that need to be mined – depends on the required physical and chemical properties of the products. in some cases more than one material may fulfil the product requirements and the choice will then be based on price and availability. all societies need mineral resources for their development, but exactly which minerals and metals are in demand and how they are used depend on the stage of development of the particular society. during historic time the trend has been very clear; innovation and new technologies require an increasing number of specialised raw materials. consequently, we need to explore for new types of minerals to meet new demands. the demand for minerals is fueled by a number of drivers demographics – the united nations has estimated that the world population will increase from currently 7 billion to 9 billion by 2050 and that about 6.5 billion people will © 2014 geus. geological survey of denmark and greenland bulletin 31, 95–98. open access: www.geus.dk/publications/bull 9696 live in cities in 2050. this trend creates a need to develop new infrastructure to support the fast-growing urbanisation, which in turn creates an increased demand for minerals, in particular sand, gravel, iron and copper. numerous other raw materials are also needed for basic infrastructure. wealth – the economic growth in some of the emerging markets – e.g. brazil, russia, india, indonesia, china, the republic of korea, south africa – creates millions of new customers for products like houses, household machines, bicycles, cars, computers, etc. these are all manufactured from raw materials which have to be mined and processed. an example of this is china that has the world’s largest population and is globally the largest consumer of copper, aluminium and iron. however, the consumption of copper in china is still only 3 kg/person/year, much lower than in europe where the consumption is 16 kg/person/year (bogner 2012). however, it is expected that china’s copper consumption will increase substantially mainly as a result of growing wealth, rather than just the growing population. technology – the introduction of new materials, for example in houses and vehicles, in new electronic communication equipment and in new ‘green’ energy technology, changes the desired physical and chemical properties of materials, which in turn creates demand for new mineral raw materials. emerging technologies and new materials have created a rapidly growing demand for certain commodities such as indium and gallium used in light-emiting diode lamps; lithium, copper, neodymium and dysprosium used in electric cars; indium, cadmium and tellurium in photovoltaic thin-film and dysprosium and neodymium in magnets. concurrently, the need for some traditional materials has been reduced. for example, light, strong materials such as aluminium and magnesium have reduced the amount of steel required to build car frames. critical minerals and vulnerability to supply restrictions during the past decade mineral resource shortage has made headlines in the media, especially with regard to the ree. in response, a number of institutions have developed lists of mineral criticality on regional and national levels. for example, the european union has defined 14 raw materials as critical to the eu (european commission 2010). typically, the studies have used a two-fold approach: (1) assessment of the supply risk and (2) assessment of the impact of an actual shortage. the term critical minerals is frequently used in this context. critical minerals are those which are important to su pp ly ris k geological, technological, and economic social and regulatory geopolitical national vulnerability to supply restriction importancesubstitutabilitysusceptibility 1 3 2 4 fig. 1. diagram of vulnerability to the supply risk and restriction (modified from graedel et al. 2012). in the diagram element 1 has a low supply risk and even if a supply shortage occurs, this will not have a great impact on society; element 4 has a high supply risk and society is vulnerable to supply restictions; element 2 possesses a high supply risk but low vulnerability to supply restrictions; and for element 3 the opposite situation occurs, the supply risk is low, but in the event of a supply risk the national vulnerability is high. 97 society and subject to a specific availability or supply risk, e.g. at the corporate, national, regional or global industry level. scarcity is the potential outcome of criticality if a supply risk is not effectively mitigated. scarcity can be a result of several factors such as political conflicts, embargos, cartels, natural disasters, sudden increases in demand, inadequate investment in new mines and processing facilities or resource depletion. resource depletion causing significant shortages of mineral commodities has not yet been documented except in the case of cryolite, but it may pose a long-term threat. based on longand medium-term supply risk graedel et al. (2012) assessed the vulnerability to supply shortage and identified three general components, namely (1) geology, technology and economy; (2) social and regulatory factors and (3) geopolitical factors. each of these were specified by six indicators, forming the ‘supply risk axis’. the ‘vulnerability to supply restrictions axis’ is composed of another set of factors such as (1) importance, (2) substitutability and (3) susceptibility specified in eight indicators (see figs 1, 2). graedel et al. (2012) suggested that vulnerability should also include the environmental impact. forecasting and creating possibilities for adequate policies the value chains for mineral raw materials include all stages of mineral exploration, mining and the processes transforming the minerals into intermediate goods applicable for manufacturing by industrial end users. however, most of the companies in the chain may be unaware of shortor long-term market constraints or opportunities. this prevents the industry itself from responding to sudden changes in demand. the exploration that targets new raw materials is therefore driven by commodity prices. globally, 2556 companies spent 20.5 billion us$ on mineral exploration in 2012, of which 49% was spent on gold, 32% on base metals and the remaining 19% on all other commodities (wilburn & stanley 2013). this illustrates that the exploration sector is decoupled from the end user demand. furthermore, there is a mismatch between the time scales of action in different parts of the value chain. industrial demand for new raw materials and markets for raw materials fluctuate on short-time scales, whereas the time needed to adjust the supply is much longer; it typically takes more than ten years to open a new mine, and sometimes even substantially longer. scrap supplies for recycling, secondary raw materials, are insufficient and usually too expensive to handle in order to bridge the gap between short-term demand and supply. individual governments and their institutions need updated assessment data on the national vulnerability to supply restrictions of mineral raw materials in order to develop and implement policies to avoid scarcity of particular critical minerals. for example, the general conditions for europe may not necessarily be accurate and relevant for the danish industrial and agricultural sectors. so far, only very limited data on vulnerability to supply restrictions are available for substitutability susceptibilityimportancecomponent indicator 87.5 (75–100) 62.5 (50–75) 37.5 (25–50) 12.5 (0–25) national economic importance net import reliance ratio net import reliance global innovation index substitute performance substitute availability environmental impact ratio see equation in si score for percentage of population utilising percentage of population utilising supply risk score of substitute see equation in si see equation in si see equation in si see equation in si poor adequate good exemplary sc or e fig. 2. components of the valuation methodology for the vulnerability to supply restriction, detailing the x-axis in fig. 1 (from graedel et al. 2012). supporting information (si) is detailed in: http://pubs.acs.org/doi/suppl/10.1021/es203534z/suppl_ file/es203534z_si_001.pdf 9898 public and private stakeholders in denmark. in 2013, the geological survey of denmark and greenland (geus) established the center for minerals and materials (mima) to identify and study the most important raw material value chains. the danish government subsequently decided to strengthen the knowledge about criticality, vulnerability and scarcity of raw materials and have requested mima to carry out a three-year research programme to complete a vulnerability analysis for denmark. mima is currently identifying an adequate approach for this programme. danish industry is characterised by an advanced downstream sector that depends on many imported components in end-product assemblages, while manufacturing of upstream products based on primary raw materials is of lesser importance. however, regardless of where the danish manufacturing activities belong in the value chains, they are all based on mineral raw materials, some of which may be classified as critical minerals. it is important to examine and map the extent to which denmark is subject to supply restrictions and to understand the implications of such vulnerability. danish consumers may not be aware of a product’s requirements with regard to raw materials, and thus remain unaware of a potential supply problem attached to the product. statistically, denmark monitors export and import of all goods in compliance with international categories for goods and industries, but there is a need for more knowledge about the amount and types of processed raw materials in these goods and components used by danish industry. mima and its partners will investigate these issues further and disseminate results, analyses and forecasts. conclusions denmark, like all other countries, depends on mineral raw materials – domestic and imported – to sustain and develop society and is thus vulnerable to mineral raw materials scarcity. however, most consumers and companies in the downstream parts of the value chains as well as decision makers in the administration and industry are relatively unaware of this. it is the aim of the center for minerals and materials, mima, to build knowledge and disseminate information for the danish society about mineral resource supply risks and vulnerability to supply restrictions. references bogner, s. 2012: the commodity megatrend. resource investor (www.resourceinvestor.com/2012/05/09/the-commodity-megatrend). european commission 2010: critical raw materials for the eu. report of the ad-hoc working group on defining critical raw materials, 84 pp. brussels: european commission. graedel, t.e. et al. 2012: methodology of metal criticallity determination. environmental science & technology 46, 1063–1070. malthus, t.r. 1798: an essay on the principle of population, 388 pp. london: j. johnson. national research council 2008: minerals, critical minerals, and the u.s. economy. washington, dc: the national academies press. rosenau-tornow, d., buchholz, p., riemann, a. & wagner, m. 2009: assessing the long-term supply risks for mineral raw materials – a combined evaluation of past and future trends. resources policy 34, 161–175. undp 2010: human development report 2010 – 20th anniversary edition. the real wealth of nations: pathways to human development, 238 pp. published for the united nations development programme. basingstoke: palgrave macmillan. wilburn, d.r. & stanley, k.a. 2013: exploration review. annual review 2012. mining engineering, may 2013, 22–42. authors’ addresses p.k., r.j.c. & k.h., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark; e-mail: pka@geus.dk n.f., department of geosciences and natural resource management, øster voldgade 10, dk-1350 copenhagen, k, denmark. http://www.resourceinvestor.com/2012/05/09 http://www.resourceinvestor.com/2012/05/09 mailto:pka@geus.dk characterising brines in deep mesozoic sandstone reservoirs, denmark research article characterising brines in deep mesozoic sandstone reservoirs, denmark   hanne d. holmslykke*1, niels h. schovsbo1, lars kristensen1, rikke weibel1 and lars henrik nielsen1 *corresponding author: hanne d. holmslykke | e-mail: hdh@geus.dk 1geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350, copenhagen k, denmark geus bulletin vol 43 | e2019430104 | published online: 17 july 2019 https://doi.org/10.34194/geusb-201943-01-04   the danish subsurface contains several sandstone units, which represent a large geothermal resource (vosgerau et al. 2016). currently, only three geothermal plants are operating in denmark, but several exploration licences are expected to be awarded in 2019. geothermal energy is exploited from deeply buried porous sandstones by bringing warm formation water (brine) to the surface, extracting the heat and returning the cooled water to the same sandstones. the reduced temperature of the brine during this process implies a risk of scaling, which may reduce reservoir permeability and hence injectivity. predicting the chemical composition of formation waters, however, could help to reduce the risk associated with scaling in planned geothermal facilities. here, we present a regional overview of the geochemistry of brines from deep mesozoic sandstones in the danish basin and north german basin that supplements previous studies, notably by laier (2002, 2008). the brine composition at shallow burial typically reflects the original (connate) formation water chemistry, which is determined by the original depositional environment of the sandstone, for example fluvial or marine. however, the mineralogical composition of the sandstone changes during burial, whereby some minerals may dissolve or precipitate when exposed to higher temperatures. these mineral changes are reflected in the brine composition, which typically becomes more saline with increased burial (e.g. laier 2008; kharaka & hanor 2003). the brine chemistry reported here shows a distinct depth trend, which reflects original connate formation waters that are modified through burial diagenesis. we have classified the brines into brine types, which are shown to be related to their depositional environment, depth, geological formation and geographical domains. methods we collected new samples from the production wells at each of the three danish geothermal sites (margretheholm, sønderborg and thisted) in 2017. the samples were analysed for ph, anions, cations and trace elements. cations were analysed by icp-ms (perkinelmer elan6100drc quadrupol) with a standard deviation of 3–15% depending on the element measured. samples for anion analysis were frozen for ion-chromatography (lc50-cd50, dionex, ca, usa) with a quantification limit of 0.05 mg/l. total dissolved species (tds) were calculated as the weight sum of analysed ions per weight of saltwater. these new data are combined with previously published data from the northern north german basin (tønder-4, -5 wells, sønderborg-1a, -2 wells) and from the danish basin (fig. 1, farsø-1, aars-1, stenlille-1, -19 wells, thisted -2, -3 wells, margretheholm-1, -2 wells; laier, 2002, 2008; hjuler et al. 2019). collectively, these samples span the triassic skagerrak, bunter sandstone and falster formations, the upper triassic – lower jurassic gassum formation and the jurassic fjerritslev and haldager sand formations (fig. 2). contaminated samples picked during test pumping were deselected and only samples in the late phase with stable water chemistry (notably k+ and cl-) were used. a principal component analysis (pca) was conducted on the combined dataset to classify the formation water chemistry into one of three brine types (table 1). altogether, the new samples and the samples from the literature total 39 samples (fig. 3). in the pca, present day vertical depth is used with no correction for cenozoic uplift. table 1.  brine type characteristics. brine type depth cl– br– so42– na+ k+ ca2+ mg2+ ca:cl tds ph sr2+ fetot mntot nh4 + zn2+ ba2+ li+ sio2 m ppm ppm ppm ppm ppm ppm ppm mol/mol ppm ppm ppm ppm ppm ppm ppm ppm ppm 1a 1250 97 264 310 869 53 382 502 4372 1167 0.04 157 866 6.6 196 19 6 0 1 2 1b 1376 105 000 372 26 56 667 421 8033 1480 0.08 171 998 6.3 498 39 15 27 11 15 10 13 3 1769 195 250 1275 595 115 000 1425 8750 1400 0.04 323 695 2b 2511 177 667 993 163 77 333 1838 27 950 2820 0.16 288 765 5.9 815 281 80 20 28 9 28 2a 2565 135 407 824 261 53 758 717 22 276 2820 0.16 216 064 6.3 862 6 20 21 5 11 11 6 tds: total dissolved species. fetot: total iron. mntot: total manganese. fig. 1.  location of the wells used for brine type characterisation and distribution of geothermal reservoirs: skagerrak (yellow – orange), bunter sandstone (red), and gassum (blue) formations. modified from weibel et al. (2017a). fig. 2.  stratigraphy of reservoirs and location of brine types. brine types are defined in fig. 3. the broken line indicates present-day depth of 2 km below surface. tø: tønder. sø: sønderborg. aa: aars. fa: farsø. ti: thisted. st: stenlille. ma: margretheholm. palaeo.: palaeozoic. cen.: cenozoic. fig. 3.  principal component analysis (pca). a: score plot. b: loading plot of pca on regional data. the plots model 85% of the total data variance. variance proportions are shown along each component axis. brine types are classified according to their groupings in a. results and discussion to classify the brine types we applied a pca, which transforms a matrix of measured data (x; comprised of n samples and p variables) into sets of projection subspaces delineated by principal components. each component is a linear combination of all p variables, which displays variance-maximised interrelationships between the variables (schovsbo et al. 2016). pca results are presented as a score plot (fig. 3a), which displays groupings, or clusters, of samples based on compositional similarities, alongside a loading plot (fig. 3b), which shows the variable correlations. finally, we quantify the proportion of the total dataset variance that can be modelled by each component (as shown in each of the axes in fig. 3). brine type classification in the pca model, the first two principal component axes resolve 85% of the total data variance (fig. 3). the main trend expressed on the pca-1 axis is high ion concentrations in water from deeply buried reservoirs, expressed as high positive pca-1 loadings at depth (fig. 3a) for all ions except so42– (fig. 3b). the pca-2 axis displays high positive loadings of na+, cl–, k+, so42– and br–, as well as high negative loadings of elements like mg2+, ca2+ and also at depth (fig. 3b). from the pca, three main brine types can be identified based on natural groupings in the pca-2 versus pca-1 plot (fig. 3a). the characteristics and occurrence of each type are presented below. brine type 1 plots with negative pca-1 and with slightly positive (type 1a) or negative (type 1b) pca-2 score (fig. 3a) and is characterised by relative low chloride content (cl– < 110 000 ppm) and low tds (table 1). type 1a is so42– enriched, and type 1b is so42– depleted (table 1). type 1 occurs mainly in gassum reservoirs from the stenlille, thisted (type 1b) and sønderborg wells (type 1a). these three reservoirs are buried between 1.2 and 1.6 km depth. brine type 2 is characterised by positive pca-1 and negative (type 2a) or neutral (type 2b) pca-2 scores and can be characterised compositionally by medium to high salinities (cl– >130 000 ppm), and high to very high ca2+ and mg2+ concentrations (fig. 3). type 2a is cl– and k+ depleted relative to 2b, whereas type 2b is k+ enriched (table 1). brine type 2 occurs at depths greater than 2 km in a broad range of reservoirs belonging to the haldager sand, gassum, bunter sandstone and skagerrak formations. type 2a occurs exclusively in the margretheholm area (fig. 3a). brine type 3 has intermediate pca-1 and highly positive pca-2 scores (fig. 3a), reflecting high salinities (cl– >190 000 ppm, table 1). type 3 waters are restricted to the bunter sandstone reservoir in the tønder area overlying a zechstein salt dome. in addition, halite (nacl) is also present in the triassic sequence at tønder (laier & nielsen 1989). the brine is halite saturated and it is estimated that c. 5 g/l halite has precipitated at surface conditions during production tests (hjuler et al. 2019). saturation index from phreeqc simulations to further interpret the brine types, we performed a chemical speciation analysis using the numerical code phreeqc and its pitzer database. the in situ reservoir temperature was estimated from the regional temperature gradient (e.g. balling et al. 1981). the fluid pressure of the reservoirs was assumed to be hydrostatic (9.79 kpa/m) assuming a water density of 1.1 g/cm3). the saturation state of the brines with respect to selected minerals is indicated by the saturation index (si) whereby positive and negative values indicate super-saturation and undersaturation, respectively. equilibrium with respect to the mineral is assumed for –0.4 ≤ si ≤ +0.4. this accounts for the uncertainties associated with the difficulties of sampling brines at high temperature and pressure, the analytical uncertainty and the application of thermodynamic equilibrium constants on mineral phases in saline systems. formation water with a si within this band is assumed to be saturated, and thus in equilibrium for this mineral. carbonate minerals are not included in table 2, due to the difficulty of correctly measuring components of the carbonate system (ph, hco3etc.) in heated pressurised samples. table 2.  saturation index (si) for selected minerals calculated with phreeqc. brine type depth (m) anhydrite halite barite celestite 1a 1250 –0.42 –0.90 0.40 0.08 1b 1376 –1.69 –0.79 –0.06 –1.08 3 1769 0.35 0.39 2b 2511 –0.29 0.03 –0.20 –0.09 2a 2565 –0.20 –0.55 0.03 0.04 minerals with a saturation index in the band –0.4 ≤ si ≤ 0.4 are assumed to be in equilibrium, within the uncertainties of analyses and thermodynamic parameters in the phreeqc database. the salinity of the brine types increases with the average depth (e.g. cl– in brine types 1a, 1b, 2b, table 1). this relationship was previously explained by the diffusion of cl– from the underlying zechstein salt deposits occurring in most of the danish basin and the north german basin (laier 2002). for the deepest reservoirs (type 2b) the brine is saturated with respect to halite (si = 0.03; table 2). even though types 2a and 2b are sampled from the same depth, type 2a has a significantly lower cl-concentration and is undersaturated with respect to halite (si = –0.55, table 2). this may be because brine type 2a, from margretheholm geothermal plant, was sampled from reservoir sandstones at the basin margin, where zechstein salt deposits are absent. brine type 3, sampled in tønder has a high salt content compared to deeper brines (types 2a and b), and the brine is saturated with respect to halite. brine type 3 is the only sample located in an area where halite-cemented sandstone intervals are present and salt deposits exist both above (röth salt) and below (zechstein salt). brine type 3 is saturated with respect to anhydrite (caso4 si = 0.35; table 2), which makes sense as anhydrite is a common mineral in the bunter sandstone formation (weibel & friis 2004). despite a high si for barite (baso4) in brine type 1a (si = 0.40, table 2), we did not observe any barite precipitation in the margretheholm geothermal plant. the high saturation index probably reflects the uncertainties associated with sampling and analysing from deep reservoirs and most likely the formation water is in equilibrium with barite in the reservoir. the ca:cl ratio (table 1) is significantly higher in the deeper types (2a and 2b) compared to the shallower brines (types 1a, 1b and 3). calcite (caco3) is common in the shallow part of the gassum formation, but ankerite (ca(mg, fe, mn)(co3)2) is more abundant in the deeper parts (weibel et al. 2017b). hence, replacement of calcite with a fe and mgrich carbonate would liberate ca2+ to the formation water. similarly, ca2+ may be liberated to the pore fluid, as dolomite (camg(co3)2) becomes more stable than calcite in the deeply buried parts of the skagerrak formation (weibel et al. 2017a). the k+ content generally increases with burial depth in the gassum formation (table 1). k+ may have been liberated by albite (naalsi3o8) replacement (albitisation) of k-feldspar (kalsi3o8), which has been documented in gassum sandstones from both the aars-1 and farsø-1 cores (weibel et al. 2017b). formation water from the bunter sandstone formation generally has a similarly high or higher k+ content as the deeply buried gassum formation. this may relate to a generally higher abundance of k-feldspar and rock fragments (weibel & friis 2004), or it may originate from kcl in the under or overlying evaporites. conclusions here, we have identified and characterised three brine types present in danish geothermal reservoirs. reservoir depth and the occurrence of salt in the subsurface layer appear to be a dominant control. brine type 1 occurs in reservoirs shallower than 2 km, whereas brine type 2 occurs in both jurassic and triassic sandstones buried to more than 2 km. among the deeply buried reservoirs, type 2a brine is only found in the margretheholm wells and is interpreted to reflect the absence of salt in the subsurface around these wells in contrast to all other analysed deep wells. brine type 3 is highly saline and occurs at less than 2 km depth in the tønder area. classification of brine types according to chemical composition, highlights variable risks in potential scale and scale types and shows that local conditions must be considered prior to any new planned geothermal facility. geological information, depth and geographical domains can serve as a rough predictive tool, and will be further refined as additional data are collected from new geothermal wells. acknowledgments this contribution is part of the project geotherm (“geothermal energy from sedimentary reservoirs – removing obstacles for large scale utilization”) (no. 6154-00011b) funded by the innovation fund denmark (ifd). we thank the two reviewers, ida fabricius and nicolas marty, for their comments, which improved the manuscript. references balling, n., kristensen, j.i., breiner, n., poulsen, k.d., rasmussen, r. & saxov, s. 1981: geothermal measurements and subsurface temperature modelling in denmark. geoskrifter 16, 173 pp. hjuler, m.l., olivarius, m., boldreel, l.o., kristensen, l., laier, t., mathiesen, a., nielsen, c.m. & nielsen, l.h. 2019: multidisciplinary approach to assess geothermal potential, tønder area, north german basin. geothermics 78, 211–223. https://doi.org/10.1016/j.geothermics.2018.12.001 kharaka, y.k. & hanor, j.s. 2003: deep fluids in the continents: i. sedimentary basins. in: drevor, j.i. (ed.): treatise on geochemistry 5, 499–540. https://doi.org/10.1016/b0080437516/05085-4 laier, t. 2002: vurdering af udfældningsrisici ved geotermisk produktion fra margretheholmboringen mah-1a. beregning af mætningsindeks for mineraler i saltvand fra danmarks dybere undergrund. danmarks og grønlands geologiske undersøgelser rapport 2002/95. 48pp laier, t. 2008: chemistry of danish saline formation waters relevant for core fluid experiments. fluid chemistry data for lab experiments related to co2 storage in deep aquifers. danmarks og grønlands geologiske undersøgelser rapport 2008/48. laier, t. & nielsen, b.l. 1989: cementing halite in triassic sandstone (tønder southwest denmark) as a result of hyperfiltration of brines. chemical geology 76, 353–363. https://doi.org/10.1016/0009-2541(89)90103-4 schovsbo, n.h., hedegaard, k., holmslykke, h.d., kjøller, c., kristensen, l., thomsen, e. & esbensen, k.h. 2016: formation water and produced water types in danish oil and gas fields: implications for enhanced oil recovery by “smart” water. geological survey of denmark and greenland bulletin 35, 43–46. vosgerau, h. et al. 2016: a webgis portal for exploration of deep geothermal energy based on geological and geophysical data. geological survey of denmark and greenland bulletin 35, 23–26. weibel, r. & friis, h. 2004: opaque minerals as keys for distinguishing oxidising and reducing diagenetic conditions in the lower triassic bunter sandstone, north german basin. sedimentary geology 169, 129–149. https://doi.org/10.1016/j.sedgeo.2004.05.004 weibel, r. et al. 2017a. the influence of climate on early and burial diagenesis of triassic and jurassic sandstones from the norwegian – danish basin. the depositional record 3, 60–91. https://doi.org/10.1002/dep2.27 weibel, r., olivarius, m., kristensen, l., friis, h., hjuler, m.l., kjøller, c., mathiesen, a. & nielsen, l.h. 2017b: predicting permeability of low enthalpy geothermal reservoirs: a case study from the upper triassic − lower jurassic gassum formation, norwegian – danish basin. geothermics 65, 135–157. https://doi.org/10.1016/j.geothermics.2016.09.003 how to cite holmslykke, h.d., schovsbo, n.h., kristensen, l., weibel, r. & nielsen, l.h. 2019: characterising brines in deep mesozoic sandstone reservoirs, denmark. geological survey of denmark and greenland bulletin 43, e2019430104. https://doi.org/10.34194/geusb-201943-01-04 e2019430302-01 the use of unmanned aerial systems (uas), also known as drones, is becoming increasingly important for geological applications. thanks to lower operational costs and ease of use, uas offer an alternative approach to aircraft-based and ground-based geoscientific measurements (colomina & molina 2014). magnetic and hyperspectral uas surveys hold particular promise for mineral exploration, and several groups have recently published studies of magnetic data collected by uas for such applications (malehmir et al. 2017; cunningham et al. 2018), although equivalent studies using hyperspectral data are still rare (kirsch et al. 2018). combining both techniques is particularly useful. magnetic measurements play an important role in mineral exploration, since magnetisation in rocks is mainly associated with magnetite and other iron minerals, which can be used in mapping and targeting of mineral deposits (dentith & mudge 2014). hyperspectral imaging (hsi) is a powerful exploration and mapping technique in areas where the rock surface is well-exposed, and where geological units and mineral compositions can be estimated from spectral features of the electromagnetic spectrum in the visual and infrared range. this paper reports on the setup of uas multi-sensor systems that can collect both magnetic and hyperspectral data, developed within the eu-funded mulsedro project (multi-sensor drones). we focus on small, lightweight solutions with take-off weights less than 5 kg. this is because regulations restrict commercial uas operations in many countries according to take-off weight, flight height and operating range – typically, only operations in the (extended) visual line of sight are allowed. lightweight systems are also particularly advantageous to support geological field campaigns as they can be deployed quickly to survey areas of interest. here, we use both a multi-copter and a fixed-wing uas as both platforms have advantages. multi-copters can fly at developing multi-sensor drones for geological mapping and mineral exploration: setup and first results from the mulsedro project björn heincke*1, robert jackisch2, ari saartenoja3, heikki salmirinne4, sönke rapp5, robert zimmermann2, markku pirttijärvi3, erik vest sörensen1, richard gloaguen2, lisa ek6, johan bergström6, arto karinen3, sara salehi1, yuleika madriz2 and maarit middleton4 geus bulletin is free to individuals and institutions in electronic form. the author(s) retain copyright over the article contents. review article | open access geus bulletin vol 43 | e2019430302 | published online: 29 july 2019 https://doi.org/10.34194/geusb-201943-03-02 uas integrated positioning system ips magnetic sensors hyperspectral camera stereo camera data acquisition & processing uas features: total system <5 kg easy attachment and detachment of sensors accurate positioning in areas with poor gps reception + integrated physical property maps produced by combining dsm, hyperspectal and magnetic images a b fig 1. in the multi-sensor drone (mulsedro) project we are developing a: uas multi-sensor systems to gather both magnetic and hyperspectral data. b: these data are integrated with digital surface models (dsms) to map physical properties on the ground. https://doi.org/10.34194/geusb-201943-03-02 e2019430302-02 low elevation and speed, allowing them to follow strongly varying topography and collect high-resolution data that are comparable to traditional ground surveys. fixed-wing uas reach higher speeds and have longer endurance, such that larger areas can be mapped in a short time – comparable to smaller helicopter surveys. in addition, we have developed a new integrated positioning system (ips) for uas use that is not based on gps. this affords accurate positioning within areas of poor gps reception, such as mining tunnels and narrow valleys. the general concept of mulsedro is summarised in fig. 1. multi-sensor uas systems the fixed-wing system is developed by the company radai oy, finland. the main characteristic of their self-constructed planes (fig. 2a) are long flight times (up to three hours), and smooth flight trajectories, which are important for the quality of the acquired magnetic and multispectral data. the multi-copter system is developed by the helmholtz institute freiberg for resource technology (hzdr-hif), using a customised uas and frames, on which different sensors can be quickly attached and detached (fig. 2b) for efficient and flexible surveying. in both platforms, integrated global navigation satellite systems (gnss) receivers and inertial measurement units measure the positions and orientations of the uas. magnetic both the fixed-wing and the multi-copter system were equipped with three-component fluxgate magnetometers, and magnetic base stations were set up to correct for the diurnal variations of the earth’s magnetic field. flight operations are performed autonomously along a predefined survey path of inand cross-lines. barometrical sensors together with a digital elevation model are used to keep the flight altitude relative to the ground as constant as possible. minimum operational altitudes for fixed-wing and multi-copter operations are typically governed by safety margins and terrain undulations, and can be as low as c. 30 m and c. 15 m, respectively. typical line spacing for the fixed-wing surveys of 20–50 m are comparable with those of small scale helicopter surveys. however, with the multi-copters, we have already performed surveys with line spacing of only 7 m, resulting in magnetic maps with resolutions comparable to groundbased magnetic studies. software developed by radai oy is especially adapted to process magnetic data from uas and includes advanced steps, e.g. equivalent layer modelling (pirttijärvi 2003) to reduce artefacts and noise, and to compute consistent magnetic fields for a constant altitude. one potentially problematic aspect of deploying magnetic sensors on multi-copters is that the electric motors can cause significant electromagnetic noise due to the proximity to the magnetic sensor (at distances <1 m). however, we observe that most noise in our magnetic surveys (using the tho-rpx8/10 multi-copter from tholeg) occurs at high frequencies and is filtered out efficiently during processing. hyperspectral imaging and photogrammetry hyperspectral data are collected with the multi-copter platform using a frame-based snapshot camera (the rikola hyperspectral imager camera from senop oy). this instrument captures images in the visible and near-infrared part of the electromagnetic spectrum (504–900 nm) with image integration times of 1–2 s. because of the significant integration time, flying at relatively high speed would cause spatial shifts of the spectral bands within a scene (e.g. assuming a typical band integration time of 15 ms, 50 bands, and a typical multi-copter speed of 5 m/s, shifts between the first and last band of a scene (i.e. the area covered by the image stack) would be >3.5 s). this means that illumination conditions hyperspectral camera gimbal batteries gnss fluxgate magnetometer tholeg tho-r-px8/10 net weight: 5.5 kg payload: 6.5 kg flight time: 20 min albatros vt3 wingspan: 2.8 m net weight: 4 kg payload: 2 kg flight speed: 15-30 m/s flight time: up to 3 h b a fig 2. the uas systems used within the project. a: custom built fixed-wing uas from radai oy equipped with a magnetometer. b: customised multicopter from hzdr-hif equipped with a hyperspectral frame camera. e2019430302-03 would significantly change for bands of one scene which again could result in inconsistent hyperspectral products for interpretation. therefore, the multi-copter maintains a steady position in the air during data acquisition such that spectral bands of the same image strongly overlap and have little illumination changes. during flight operation, hyperspectral data are usually collected using a setting of 50 image bands with a spectral resolution of c. 8 nm and operation heights in the range of 50–100 m, which results in ground pixels resolutions of c. 3–7 cm. to drape hyperspectral images on a precise surface topography, we performed an additional uas survey to determine a high-resolution digital surface model (dsm) that covers the entire study area. a commercial fixed-wing equipped with a rgb camera is used for this survey and the dsm is determined by structure-from-motion photogrammetry using a flight pattern of parallel lines. we used the mephysto toolbox (jakob et al. 2017) to pre-process the hsi data and for integration with the dsm, which includes lens correction, co-registration (correcting for spatial mismatches of spectral bands), orthorectification (matching the hyperspectral images with orthophotos – aerial photographs, geometrically corrected such that its scale is undistorted – obtained by photogrammetry), topographic correction (reducing the influence of topography onto the local illumination), and merging of images to a hsi mosaic and conversion from radiance to reflectance. one limit of the current setup is that the spectral range of available lightweight hyperspectral sensors (<2 kg) mountable on uas is restricted to the visible near infrared. however, diagnostic absorption features of most minerals are located at higher wavelengths, in the short-, midand long-wave infrared parts of the spectrum. therefore, the mulsedro setup is currently limited to mapping iron-bearing minerals such as hematite, jarosite and goethite that have characteristic spectral features in the visible near infrared. the system can also identify vegetation features that can be used as proxies for mineralisation (i.e. different species of vegetation that correlate with the underlying rock types; gupta et al. 2018). other potential applications include mapping of rare-earth elements and monitoring mine tailings (jackisch et al. 2018). integrated positioning system the ips for uas is developed from the handheld pilot3d device (fig. 3a) for underground mine navigation and 3ddocumentation. the ips is based on a multi-sensor approach, where an inertial measurements unit (a device typically used to manoeuvre aircraft) is combined with a stereo camera (benecke et al. 2016). both sensors independently acquire positioning information and combining these data from the devices provides accurate real-time positions with accuracies of c. 0.1% of the distance covered at walking speed (c. 1–2 m/s). we have completed initial tests from drones (figs 3b, c) to evaluate the performance of the ips on faster moving platforms. application and outlook the mulsedro uas systems were tested during field campaigns in two mining areas of central finland (fig. 4a) in the summers 2017 and 2018. here, we present some of the initial results from one of these sites located in the northern part of the siilinjärvi carbonatite-glimmerite complex (63°08 4́4́ ´n, 27°44́ 16́ ´e), 1 km north of a large open pit phosphate mine. the complex intruded into granite gneiss (2610 ± 4 ma) and covers an area c. 16 × 1.5 km (o’brien et al. 2015). our objectives were to (1) trace the main subvertical carbonatite veins in areas covered by overburden by magnetic surveys, (2) identify relationships between those stereo cameras rgb camera white illumination a b fig 3. the integrated positioning system pilot3d as a: handheld device. b: mounted on a uas. c: images obtained by the stereo camera during a uas test flight. e2019430302-04 veins and the main mineralogy of the area, and to (3) identify possible indications of rare-earth elements in excavated test outcrops by hsi. we conducted fixed-wing magnetic and photogrammetry surveys (figs 4b, c) across areas of c. 1 km2. these surveys provide a larger-scaled picture of the topography and the magnetic anomalies from areas that are predominately covered by soil and vegetation. on accessible outcrops located within these areas, multi-copters equipped with hyperspectral and magnetic sensors collected data with significantly higher resolution (fig. 4d) – down to a few centimetres for hyperspectral data. to connect these magnetic anomalies and spectral characteristics with rock properties on the ground, the uas measurements were supplemented with ground-based investigations, including handheld xrf, spectroradiometry, magnetic susceptibility measurements and collection of rock samples for analysis in the laboratory. in this way, our survey strategy comprises the full range of spatial scales, from kilometre to sub-centimetre. the total magnetic field map obtained from fixed-wing uas data in siilinjärvi shows a north-south trending positive magnetic anomaly in the central part of the survey area, and smaller anomalies in the east (fig. 4c). the central anomaly coincides well with structural features from the regional geological map (see “bedrock of finland”; https:// gtkdata.gtk.fi/kalliopera/index.html) and is described as a major fault zone with strong deformation trends (salo 2016). o’brien  et  al. (2015) described the zone as an interface of glimmerite-carbonatite – present as fine and elongated carbonatite lenses in a phlogopite matrix – and surrounding b c d e a water debris fenite carbonatite carbonatite fenite diabase fig 4. a: location of field test sites in finland. note that only results from siilinjärvi site, but not from otanmäki site are presented here. b: orthophoto from photogrammetry. c: total magnetic intensity maps derived from fixedwing uass surveys of the siilinjärvi test site. d: image from a processed hyperspectral multicopter-based data set draped on the digital surface model from photogrammetry. the hyperspectral image shows the combination of bands 3, 2, 1 of a minimum noise fraction transformation in rgb, where green shading is associated with mainly fenite-hosting rocks. location of the hyperspectral survey is outlined in b and c with a black rectangle. (a second survey was performed in the southern end of the field site (see b and c), but these data are not presented here). e: ground sampling for one of the outcrops covered with hyperspectral surveying. https://gtkdata.gtk.fi/kalliopera/index.html https://gtkdata.gtk.fi/kalliopera/index.html e2019430302-05 fenites. although carbonatites and fenites at the open test outcrops did not show diagnostic spectral absorptions in the spectral range of the hyperspectral camera used here, separation of the hsi mosaic in the minimum noise fraction plot is indicative of the different lithologies (fig. 4d). field validation was achieved by finding characteristic hsi features that are associated with carbonatite bearing units (fig. 4d) from ground-based spectroradiometry, thin-section microscopy and mineralogy and geochemistry from rock samples (unpublished data, robert jackisch). future work will investigate how information from all datasets – both field validation and in particular from the uas surveys – can be integrated to gain an improved understanding of the relationships between magnetic and hyperspectral properties and rock composition and, hence, surface geology and mineral deposit characteristics. we plan to deploy our uas based systems to an exploration target in the remote, arctic environment of greenland in the summer of 2019. finally, we will integrate a multispectral sensor on a fixed-wing uas to capture multispectral information over larger areas of several kilometres within a single flight. acknowledgements the mulsedro project is funded by eit rawmaterials (project number: 16193). we thank aleksi salo (yara suomi oy) for allowing us to undertake surveys within yara’s premise. comments from the reviewers bo møller stensgaard and alireza malehmir improved the manuscript. references benecke, n., born, a., boerner, a., rapp, s., stelzer, p., tsirigotis, n., weber, m. & zuev, s. 2016: mobile solution for positioning, 3d-mapping and inspection in underground mining. 16th international congress for mine surveying, brisbane, australia. 12–16 sept. 2016. colomina, i. & molina, p. 2014: unmanned aerial systems for photogrammetry and remote sensing: a review. isprs journal of photogrammetry and remote sensing 92, 79–97. https://doi.org/10.1016/j. isprsjprs.2014.02.013 cunnigham, m., samon, c., wood, a. & cook, i. 2018: aeromagnetic surveying with a rotary-wing unmanned aircraft system: a case study from a zinc deposit in nash creek, new brunswick, canada. pure applied geophysics 175, 3145–3158. https://doi.org/10.1007/s00024017-1736-2 dentith, m. & mudge, s.t. 2014: geophysics for the mineral exploration geoscientist. cambridge university press. https://doi.org/10.1017/ cbo9781139024358 gupta, r.p. 2018: geobotanical guides. in: remote sensing geology. third edition, springer-verlag, berlin, 347–348. https://doi. org/10.1007/978-3-662-55876-8_19 jackisch, r., lorenz, s., zimmermann, r., möckel, r. & gloaguen, r. 2018: drone-borne hyperspectral monitoring of acid mine drainage: an example from the sokolov lignite district. remote sensing 10, 385. https://doi.org/10.3390/rs10030385 jakob, s., zimmermann, r. & gloaguen, r. 2017: the need for accurate geometric and radiometric corrections of drone-borne hyperspectral data for mineral exploration: mephysto-a toolbox for pre-processing drone-borne hyperspectral data. remote sensing 9, 88. https://doi. org/10.3390/rs9010088 kirsch, m., lorenz, s., zimmermann, r., tusa, l., möckel, r., hödl, p., booysen, r., khodadadzadeh, m. & gloaguen, r. 2018: integration of terrestrial and drone-borne hyperspectral and photogrammetric sensing methods for exploration mapping and mining monitoring. remote sensing 10, 1366. https://doi.org/10.3390/rs10091366 malehmir, a., dynesius, l., paulusson, k., paulusson, a., johansson, h., bastani, m., wedmark, m. & marsden, p. 2017: the potential of rotary-wing uav-based magnetic surveys for mineral exploration: a case study from central sweden. the leading edge, 552–557. https://doi. org/10.1190/tle36070552.1 salo, a. 2016. geology of the jaakonlampi area in the siilinjärvi carbonatite complex. 30 pp. bachelor’s thesis. university of oulu, finland. o’brien, h., heilimo, e. & heino, p. 2015: the archean siilinjärvi carbonatite complex. in: maier, w., o’brien, h. & lahtinen, r. (eds): mineral deposits of finland, 327–343. elsevier. https://doi. org/10.1016/b978-0-12-410438-9.00013-3 pirttijärvi m. 2003: numerical modelling and inversion of geophysical electromagnetic measurements using a thin plate model. phd thesis, university of oulu, finland. how to cite heincke, b., jackisch, r., saartenoja, a., salmirinne, h., rapp, s., zimmermann, r., pirttijärvi, m., sörensen, e.v., gloaguen, r., ek, l., bergström, j., karinen, a., salehi1, s., madriz, y., middleton, m. 2019: developing multi-sensor drones for geological mapping and mineral exploration: setup and first results from the muledro project. geological survey of denmark and greenland bulletin 43, e2019430302. https://doi.org/10.34194/geusb-201943-03-02 *corresponding author: björn heincke | e-mail: bhm@geus.dk 1 department of glaciology and climate, geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350, copenhagen k, denmark. 2 helmholtz-zentrum dresden-rossendorf, helmholtz institute freiberg for resource technology, chemnitzer str. 40, 09599 freiberg, germany 3 radai oy, teknologiantie 18, 90590 oulu, finland 4 geological survey of finland, lähteentie 2, fi-96101 rovaniemi, finland 5 deutsche montan technologie gmbh, am technologiepark 1, 45307 essen, germany 6 ltu business ab, aurorum 1, 97775 luleå, sweden https://doi.org/10.1016/j.isprsjprs.2014.02.013 https://doi.org/10.1016/j.isprsjprs.2014.02.013 https://doi.org/10.1007/s00024-017-1736-2 https://doi.org/10.1007/s00024-017-1736-2 https://doi.org/10.1017/cbo9781139024358 https://doi.org/10.1017/cbo9781139024358 https://doi.org/10.1007/978-3-662-55876-8_19 https://doi.org/10.1007/978-3-662-55876-8_19 https://doi.org/10.3390/rs10030385 https://doi.org/10.3390/rs9010088 https://doi.org/10.3390/rs9010088 https://doi.org/10.3390/rs10091366 https://doi.org/10.1190/tle36070552.1 https://doi.org/10.1190/tle36070552.1 https://doi.org/10.1016/b978-0-12-410438-9.00013-3 https://doi.org/10.1016/b978-0-12-410438-9.00013-3 https://doi.org/10.34194/geusb-201943-03-02 mailto:bhm%40geus.dk?subject= geological survey of denmark and greenland bulletin 41, 2018, 29-32 29 in recent years there has been an increased interest in neogene hydrocarbon accumulations in the north sea. the production of gas from pliocene–quaternary deposits in the dutch sector, the discovery of oil-bearing miocene sands in the lille john area and oil accumulation in middle miocene deposits in the t-1 well in the northern part of the danish central graben area, have documented neogene hydrocarbon accumulations. some of these deposits are of economic interest. this study presents an oil-bearing, middle miocene diatom ooze in the valhall field (well 2/8-g-2), within the norwegian sector (fig. 1). the valhall field is located just north of the danish–norwegian sector boundary. geological setting during the miocene, the north sea formed a silled-bounded basin with connection to the atlantic ocean via a strait between norway and shetland (rasmussen et al. 2008; fig. 1). the main sediment source areas were the shetland platform, which supplied sediments to the northern north sea and the southern scandes from which sediments were routed southwards into the south-eastern north sea (fig. 1). during the early miocene relatively large delta complexes formed from these areas and resulted in eastward progradation off the shetland platform (skade formation; eidvin et al. 2014) and south-westward progradation south of the southern scandes (ribe group; rasmussen et al. 2010). during the miocene oil-bearing diatom ooze from the north sea emma sheldon, erik s. rasmussen, karen dybkjær, tor eidvin, fridtjof riis and rikke weibel 0°e 10°e 20°e 50°n 60°n shetland platform so ut he rn sc an de s 2/8-g-2 mittelgebirge highs fig. 1. palaeogeographical reconstruction of the early miocene north sea. note that the main sediment influx from the shetland platform (yellow arrows) filled the northern north sea and that sediment supply to the eastern north sea had its source in southern scandes. based on rasmussen et al. (2008). b c a 10 µm 20 µm 2 mm fig. 2. cored diatom ooze from the valhall field, norwegian sector of the north sea. a: optical microscope image of chip of diatom ooze, note the diatom in the upper part (black arrow). b: ?denticulopsis kanayae. c: ?denticulopsis nicobarica. © 2018 geus. geological survey of denmark and greenland bulletin 41, 29–32. open access: www.geus.dk/bulletin http://www.geus.dk/bulletin 3030 middle miocene these delta complexes where flooded due to reorganisation of the tectonic regime in north-west europe which resulted in accelerated subsidence of the basin margins. consequently, much of the north sea was starved of sediment during the middle and early late miocene. the north sea area was located in the western wind belt with seasonal storms. therefore, the coast was strongly influenced by wave action. in the deeper basin, which was up to c. 1000 m deep, hemipelagic deposition predominated. a counter-clockwise current system redistributed and reshaped muddy sediments along the delta and shelf slopes within the basin (e.g. hansen et al. 2004). during the early miocene a humid, warm temperate climate predominated, similar to present day western florida (usa). a change to a cooler climate commenced in the middle miocene which probably also resulted in the enhanced influence of cold-water current systems from the atlantic ocean. under these cooler climatic conditions diatoms bloomed and resulted in deposition of diatom ooze. diatom ooze the diatom ooze is fine-grained and grey to brown, since it contains oil (fig. 2). diatom valves and radiolarians could be seen in optical and scanning electron microscopes (figs 3, 4). the diatom ooze consists of a mixture of abundant diatom valves, radiolarians and clay (fig. 4). the porosity is 50–60 %. porosity is mainly associated with diatom valves and either occurs inside the valves or in pockets next to the valves or other fossils. reduced porosity is observed in some samples, attributed to clay intruding into the diatom valves. the permeability is assumed to be low, due to the small size of the pores and tortuous connectivity between the largest pores. the measured porosity of the diatom ooze would correspond to a permeability of 0.006–0.02 md in diatom ooze from various localities in the pacific ocean (gamage et al. 2011). nannofossils and microfossils core sample (1802.7 m from the well 2/8-g-2) was analysed for nannofossil, microfossil and the presence of diatoms. diatom valves and debris were found to be common. diatoms include ?thalassiosira spp., ?denticulopsis kanayae and ?denticulopsis niobarica. d. kanayae and d. nicobarica range from the early to middle miocene (barron 1985; fig. 2). the sample was barren with respect to nannofossils and the mi2 µm 10 µm 10 µm50 µm fig. 3. four scanning electron microscope images of the diatom ooze. note the valve of the centric diatom thalassiosira spp. in the upper right image. 31 crofossil fraction yielded one radiolarian (cenodiscus spp.) and no foraminifera. palynology six core samples were analysed for palynology (1797.0 m, 1802.7 m, 1803.4 m, 1813.0 m, 1819.0 m and 1827.0 m). in all samples, the assemblages of organic particles are characterised by a dominance of marine dinoflagellate cysts (dinocysts). bisaccate and non-saccate pollen and wood particles occur very sporadically while no freshwater algae were recorded. the dinocyst assemblage is rich and diverse and the consistent presence of nematosphaeropsis spp. and impagidinium spp. indicates an outer neritic to oceanic setting (brinkhuis 1994). an increase in cold-water tolerant dinocyst taxa (mainly habibacysta tectata) was found (fig. 5), ranging from no recordings in the lowermost sample, to sporadic occurrences in the next samples and common occurrences in the two uppermost samples. the occurrences of the dinocyst species unipontodinium aquaductum in all six samples (fig. 5) strongly indicate that the cored interval should be referred to the unipontodinium aquaductum zone of dybkjær and piasecki (2010). the age of this zone is mid-langhian to early serravallian (middle miocene). the a b c d 1 2 3 4 5 1) stephanopyxis turris (diatom) 2) cross section of a diatom valve 3) ? peridinium longispinum (radiolarian) 4) ? stephanopyxis turris 5) ????? 30 µm 30 µm 30 µm 30 µm fig. 4. optical microscope images of the diatom ooze. the ooze comprises siliceous microfossils (e.g. diatoms and radiolarians) and clay. porosity inside microfossils is recognised by the blue staining of the epoxy impregnating the ooze. a and b are identical; b with crossed nicols. 1: stephanopyxis turris (diatom). 2: cross section of a diatom valve. 3: ?peridinium longispinum (radiolarian), 4: ?stephanopyxis turris. 5: diatom frustule. alexander mitlehner, uk, kindly helped with the identification of the diatoms. 20 µm 20 µm a b fig. 5. dinoflagelate cysts from the oil-bearing cores. a: unipotodinium aquaductum. b: habibacysta tectata. increased abundance of habibacysta tectata, a cold-water tolorant species, is probably associated with the climatic deterioration in the middle miocene (serravalian). 3232 unipontodinium aquaductum zone occurs in the upper part of the hodde formation defined onshore denmark which correlates with the lowermost part of the nordland group. depositional environment based on palynology and seismic stratigraphic studies (e.g. rasmussen et al. 2005), the depositional setting was outer neritic to oceanic, with a water depth just below 1000 m. late early miocene biosiliceous, organic-rich sediments of the upper lark formation in the central graben area have been described by sulsbrück & toft (2018). this part of the upper lark formation was deposited at the termination of shoreline progradation from the southern scandes (ribe group). the studied section represents slightly younger deposits than those laid down during the transgression of the lower miocene ribe group. consequently, the depositional environment was sediment starved and fully marine. the occurrence of cold water dinocysts in the studied cores, probably reflects the beginning of the middle miocene (serravalian) global climatic deterioration (zachos et al. 2001). petroleum system the oil-bearing miocene diatom ooze from the norwegian sector of the north sea described here documents oil migration into younger deposits, which are normally considered to be non-prospective. in the danish north sea area, a number of wells have penetrated hydrocarbon-bearing strata of miocene and pliocene ages as described above. the oil-bearing deposits are found in the western and central parts of the danish and norwegian central graben. the oil probably has a source in the jurassic shale deep in the central graben. migration into cenozoic deposits probably occurred along salt structures. due to early quaternary tilting of the north sea basin (rasmussen et al. 2005), up-dip migration into stratigraphic and structural traps located in the eastern part of the central graben area and the ringkøbing–fyn high may have occurred. this calls for a total re-evaluation of the petroleum system of the cenozoic succession in the north sea area. references barron, j.a. 1985: miocene to holocene planktic diatoms. in: bolli, h.m., saunders j.b. & perch-nielsen, k. (eds): plankton stratigraphy, 763–809. cambridge: cambridge university press. brinkhuis, h. 1994: late eocene to early oligocene dinoflagellate cysts from the priabonian type-area (northeast italy): biostratigraphy and paleoenvironmental interpretation. palaeogeography, palaeoclimatology, palaeoecology 107, 121–163. dybkjær, k. & piasecki, s. 2010: neogene dinocyst zonation in the eastern north sea basin, denmark. review of palaeobotany and palynology 161, 1–29. eidvin, t., riis, f. & rasmussen e.s. 2014: oligocene to lower pliocene deposits of the norwegian continental shelf, with correlation to the norwegian sea, greenland, svalbard, denmark and their relation to the uplift of fennoscandia. marine and petroleum geology 56, 184– 221. gamage, k., screaton, e., bekins, b. & aiella, i. 2011: permeability-porosity relationships of subduction zone sediments. marine geology 279, 19–36. hansen, j.p.v., clausen, o.r. & huuse, m. 2004: 3d seismic analysis reveals the origin of ambiguous erosional features at a major sequence boundary in the eastern north sea: near top oligocene. geological society memoirs (london) 29, 83–90. rasmussen, e.s., vejbæk, o.v., bidstrup, t., piasecki, s & dybkjær, k. 2005: late cenozoic depositional history of the danish north sea basin: implications for the petroleum systems in the kraka, halfdan, siri and nini fields. in: dore, a.g. & vinding, b.a. (eds): petroleum geology: north-west europe and global perspectives. proceedings of the 6th petroleum geology conference, 1347–1358. london: geological society. rasmussen, e.s., heilmann-clausen, c., waagstein, r. & eidvin, t. 2008: tertiary of norden. episodes 31, 66–72. rasmussen, e.s., dybkjær, k, & piasecki, s. 2010: lithostratigraphy of the upper oligocene – miocene succession of denmark. bulletin of the geological survey of denmark and greenland 22, 92 pp. sulsbrück, h. & toft, j. 2018: a new observation of a biosiliceous opal bearing sequence in the miocene lark formation in the danish north sea. 33rd nordic geological winter meeting, lyngby. abstract http://2dgf.dk/foreningen/33rd-nordic-geological-winter-meeting/ ngwm-2018-abstracts/3-sedimentary-rocks-and-processes/ zachos, j.c., pagani, m., sloan, l., thomas, e. & billups, k. 2001: trends, rhythms, and aberrations in global climate 65 ma to present. science 292, 686–693. authors’ addresses e.s., e.s.r., k.d. & r.w., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: es@geus.dk. t.e. & f.r. norwegian petroleum directorate (npd), p. o. box 600, n-4003 stavanger, norway. http://2dgf.dk/foreningen/33rd-nordic-geological-winter-meeting/ngwm-2018-abstracts/3-sedimentary-rocks-and-processes/ http://2dgf.dk/foreningen/33rd-nordic-geological-winter-meeting/ngwm-2018-abstracts/3-sedimentary-rocks-and-processes/ mailto:es@geus.dk geological survey of denmark and greenland bulletin 1, 543-554 543 an offshore transgressive–regressive mudstone-dominated succession from the sinemurian of skåne, sweden nils frandsen and finn surlyk a sinemurian mudstone-dominated succession was exposed until recently in the gantofta quarry in skåne, southern sweden. the deposits are placed in the döshult and pankarp members of the sinemurian–aalenian rya formation. similar facies of the same age are widespread in the danish basin where they constitute the f-ib unit (f-i member) of the fjerritslev formation. the gantofta succession thus represents the easternmost extension of the environment characteristic of the fjerritslev formation and is essentially the only locality where it has been possible to study the facies of this formation in outcrop. sedimentation seems to have taken place under relatively quiet tectonic conditions except for the possible fault-control of the basin margin. the lower part of the gantofta section is of early and early late sinemurian age. it represents the upper part of the döshult member and consists of muddy, lower shoreface sandstones, abruptly overlain by dark, bioturbated, fossiliferous mudstones with thin storm siltstones and sandstones. they are overlain by the upper sinemurian pankarp member which comprises red-brown, restricted marine calcareous mudstones with an upwards increasing number of storm siltstones and sandstones reflecting general shallowing and shoreline progradation. the succession spans the greater part of two simple sequences with a distal sequence boundary located at the boundary between the döshult member and the pankarp member. the exposed part of the lower sequence includes a thick transgressive systems tract and a very thin highstand systems tract. the upper sequence is represented by an undifferentiated transgressive and highstand systems tract. an early sinemurian sea-level rise, a late early sinemurian highstand, an early late sinemurian fall and a late sinemurian minor rise and a major fall are recognised. nearby boreholes show evidence for an end-sinemurian – early pliensbachian major rise. this evolution corresponds well with trends recorded in the subsurface fjerritslev formation of the danish basin. comparison with published european and british jurassic sea-level curves show similar overall trends, but exhibit differences in the precise ages of sequence boundaries and maximum flooding surfaces. this may reflect poor biostratigraphical resolution of the gantofta section, differences in sequence stratigraphic interpretation, real differences in the age of sequence stratigraphic key surfaces, or the basin marginal position of gantofta in the fennoscandian border zone. keywords: skåne, southern sweden, lower jurassic, sinemurian, facies analysis, sequence stratigraphy, sedimentary environments, sea-level change n.f., dong, agern allé 24–26, dk-2970 hørsholm, denmark. e-mail: nfr@dong.dk f.s., geological institute, university of copenhagen, geocenter copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. geological survey of denmark and greenland bulletin 1, 543–554 (2003) © geus, 2003 544 jurassic sedimentary rocks occur in great thicknesses in the subsurface of denmark and southern sweden (michelsen 1978; norling et al. 1983; nielsen 2003, this volume). they are, however, only exposed in a few relatively small outcrops on bornholm in the baltic sea, and in skåne, southern sweden, in the fennoscandian border zone (fig. 1; sellwood 1972; rolle et al. 1979; gravesen et al. 1982; norling et al. 1983; surlyk & noenygaard 1986; ahlberg et al. 2003, this volume). this intensely block-faulted zone forms the north-eastern boundary of the danish basin. detailed facies, biostratigraphic and sequence stratigraphic studies of exposed units are thus of outstanding importance in obtaining a more detailed picture of the sedimentary environments prevailing during jurassic time in southern scandinavia. the aims of the present paper are to interpret the facies and sequence stratigraphy of the sinemurian offshore marine deposits exposed in the gantofta quarry in skåne, southern sweden, to place the succession in its regional context, and to compare the derived sealevel curve with the sea-level curves of haq et al. (1988), hallam (1988) and hesselbo & jenkyns (1998). geological setting and stratigraphy gantofta is located close to the nw–se-trending western margin of the fennoscandian border zone (fig. 1). the margin is characterised by a major faulted flexure formed by late cretaceous – palaeogene tectonic inversion which marks the transition to the major depocentre of the danish basin to the south-west (norling 1981, fig. 37). the locality is a small clay pit (150 x 100 m) which was abandoned some years ago and the section is thus no longer easily accessible. the ammonites from the succession were described by reyment (1969a, b), the foraminifera by norling (1972), the palynology by lund (1977), the ostracodes by sivhed (1977, 1980, 1981) and the sedimentary facies and environments by frandsen (1977), rolle et al. (1979) and pieńkowski (1991a, b). the succession is 70 m thick and consists of sandstones and mudstones of the lower – lower upper sinemurian döshult member and the upper sinemurian pankarp member, which constitute the two lower members of the sinemurian–aalenian rya formation (sivhed 1984; ahlberg et al. 2003, this volume). the whole succession is tilted, and the strata strike 140° and dip 30° sw (fig. 2). vomb trough fyledalen fault gantofta 25 km höganäs basin malmö trough hanö bay kullen i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i iii i i i i i i i i i i i i i i i sweden dk bornholm i i i rhaetian and jurassic present distribution fault active in the jurassic post-jurassic fault n norway ängelholm basin fig. 1. map showing the location of gantofta in the fennoscandian border zone of southern skåne. important jurassic structural features are indicated; note that jurassic normal faults were subsequently inverted in the late cretaceous – palaeogene. based on norling & bergström (1987). dk, denmark. 545 a foraminifer fauna from the middle part of the döshult member at the base of the exposure indicates an early sinemurian age (norling 1972). an ammonite fauna including asteroceras obtusum (sowerby) and promicroceras planicostatum (sowerby) from a level 2 m below the base of the overlying pankarp member is of early late sinemurian obtusum chronozone, planicostatum subzone age (reyment 1969a). k. hoffman (in: bölau 1959) reported the occurrence of the birchi subzone (top of the turneri chronozone) immediately below the base of the pankarp member. the ostracode faunas of the döshult member indicate a latest early sinemurian to late sinemurian age (sivhed 1977). norling (1972) suggested that the succession spans the time interval of the semicostatum to obtusum chronozones. a new find of the ammonite euagassiceras cf. lundgreni reyment in the basal muddy sandstone of the section suggests a mid early sinemurian semicostatum chronozone age for this level (probably resupinatum subzone). this is in agreement with bölau (1973). correlative strata are widely distributed in the danish basin where they form the f-ib unit (f-i member) of the thickly developed, uniform mudstone package of the fjerritslev formation (michelsen 1975, 1978, 1989; pedersen 1985, 1986; michelsen et al. 2003, this volume; nielsen 2003, this volume). the succession of the gantofta quarry consists of similar facies and thus essentially represents the only locality where it has been possible to study the characteristic facies of the otherwise deeply buried fjerritslev formation in outcrop. sedimentology a detailed sedimentological log of the succession was measured in 1975–1976 by frandsen (1977). special emphasis was placed on recording primary sedimentary structures, body and trace fossils, and details of concretions and other diagenetic features were also noted. five sedimentary facies are recognised and are described below followed by an interpretation of the depositional processes and environments. sedimentary facies muddy sandstone (facies 1) this facies consists of fineto very fine-grained quartz sandstone with a mud-rich matrix. it is only known fault n a b ba i i i i i i i i i i i i i i i i i i katslösa mb pankarp mb coal seam mudstone sandstone gantofta quarry fault i i 200 m upper sinemurian lower sinemurian 200 m döshult mb fig. 2. geological map and section (a–b) of the gantofta area. modified from sivhed (1981). 546 se m ico st at um c hr on oz on e (? re su pi na tu m s ub zo ne ) ob tu su m c hr on oz on e (p la ni co st at um s ub zo ne ) trace fossils chondrites isp. rhizocorallium isp. diplocraterion isp. planolites isp. skolithos isp. pyritic tube teichichnus isp. bioturbation increasing density body fossils bivalves, suspension feeders bivalves, deposit feeders bivalve fragments gastropods ammonites belemnites scaphopods serpulids brachiopods echinoid spines shark teeth ostracodes foraminifers sequence stratigraphy sequence boundary maximum flooding surface transgressive surface of erosion transgressive systems tract highstand systems tract parasequence sb mfs tse tst hst ps tse tst mfs hst tst hst d ös hu lt m b pa nk ar p m b d ös hu lt m b lithology mudstone calcareous mudstone silty and sandy mudstone siltand sandstone carbonate cemented conglomerate structures parallel lamination parallel lamination with siltstone lenses cone-in-cone structures clay ironstone concretions c la y si lt sa nd vf f c la y si lt sa nd vf f 35 30 25 20 15 10 5 0 1 3 2 2 2 2&3 2&3 2&3 2&3 2 2 2 2 2&3 3 ps ps ps ps ps ps ps 70m m 65 60 55 50 45 40 35 2 4 5 sb fa ci es se q. st ra t. fig. 3. sedimentological section showing the biostratigraphy, chronostratigraphy and sequence stratigraphy of the sinemurian succession at gantofta. from the base of the section (fig. 3). shelly coquinas with quartz granules occur at some levels. the colour varies from light grey to dark grey. the sandstone is almost totally bioturbated and the original structures are commonly difficult to recognise. however, fine biogenic lamination caused by high densities of the trace fossil teichichnus isp. is a characteristic feature of the facies. siderite is the dominant cement, whereas the most indurated beds have a calcitic cement. the trace fossils include teichichnus isp., diplocraterion isp., rhizocorallium isp., chondrites isp., planolites isp. and zapfella isp.; the last of these occurs as elongate borings in shells of the bivalve liogryphaea arcuata. the facies contains a fully marine fauna of body fossils which occur in rather high densities. frandsen (1977) compiled a list of the fauna and documented the presence of species described from other localities by troedsson (1951). a taxonomic revision has not been attempted. the fauna includes the bivalves liogryphaea arcuata, chlamys textoria, chlamys interpunctata, entolium sp., and oxytoma sinemuriensis, the ammonite euagassiceras cf. lundgreni reyment, and indeterminate belemnites, echinoids, serpulids, ostracodes and nodosariid foraminifera. coalified wood occurs as scattered pieces up to 6 cm long. the preservation of the shells is quite variable. some are well-preserved and unworn, while others occur in coquinas and have clearly undergone some transport and destruction. the combination of marine body and trace fossils, intense bioturbation, shelly coquinas and a sand-dominated grain size indicates deposition under well-oxygenated marine conditions with normal salinity and relatively low sedimentation rates, periodically interrupted by higher energy events resulting in erosion, reworking and transport of shells. the taphonomic conditions suggest that the fauna can be considered a neighbourhood assemblage representing a fauna which lived in the area and which underwent only limited transport. the muddy nature of the sandstone suggests that the original facies was a sand-dominated, possibly flaserbedded heterolith, but the very high degree of bioturbation does not allow a detailed process interpretation. dark grey mudstone (facies 2) the mudstone of this facies characterises the bulk of the exposed part of the döshult member (fig. 3). it has a high content of silt and fine sand. silt and clay are roughly equally abundant, and the clay is dominated by kaolinite with some illite and chlorite. coaly detritus, muscovite, very small shells and shell fragments, and framboidal pyrite nodules (0.1–0.3 mm in diameter) are characteristic constituents. a few intervals, up to 1 m thick, have a relatively higher content of sand and can be classified as muddy sandstones. the mudstone is laminated with light coloured laminae of coarse silt, 1 mm thick. clay ironstone is a characteristic component and occurs as bedding-parallel siderite impregnated layers, 5 cm thick, with ellipsoidal concretions, 5–20 cm long. thin conglomerate beds consisting of reworked clay-ironstone concretions are found at the 35.05 m and 38 m levels (fig. 3). carbonate concretions with cone-in-cone structures occur at several levels. the facies is strongly bioturbated, especially in the sandier portions, but several recognisable trace fossils were noted, including diplocraterion isp., skolithos isp., rhizocorallium isp. (which has only been recorded from the clay-ironstone conglomerate) and chondrites isp. pyritic tubes, 0.2–1.0 mm in diameter, probably representing burrows of small deposit feeders, and ?planolites isp. occur throughout. this facies and the laminated siltstone–sandstone facies (facies 3) contain a rich shelly fauna. the two facies and their faunas are closely related and their faunas are described together here. the carbonate shells have undergone dissolution and are mainly poorly preserved, but wear due to transport appears to be negligible. liogryphaea arcuata is relatively rare and the specimens are smaller than those of facies 1. two species of chlamys and minute specimens of oxytoma sinemuriensis have been found. a major difference in faunal composition compared to facies 1 is the abundance of deposit-feeding bivalves of the nuculanacea (nuculana, palaeoneilo, rollieria, leda) and nuculacea (nucula). bivalves belonging to cardinia, astarte, homomya and pleuromya or related genera also occur, but the determinations are uncertain. small, high-spired gastropods representing a number of different genera are very common. poorly preserved scaphopods, brachiopods, ammonites and rare shark teeth occur at several levels. the fine-grained muds were deposited from suspension in an offshore open marine environment. the content of silt and sand probably represents material transported to the area during storms. bioturbation then resulted in mixing of the fine and coarser fractions, and destruction of primary current-produced structures. the high density and diversity of body and trace fossils show that the water was of normal salinity and well547 oxygenated. the abundance of deposit feeders is a common characteristic of fine-grained, nutrient-rich sediments, whereas the abundant small-sized gastropods suggest the presence of a marine vegetation. the clay ironstone layers were formed under conditions of negative eh, low concentration of sulphide ions, high activity of ferrous ions, and the presence of bicarbonate ions. following sellwood (1971) it is suggested that iron was transported to the marine environment and deposited as insoluble ferrioxide which was adsorbed on clay minerals. the siderite nodules contain undeformed trace fossils and the mudstone shows compaction features around the nodules which were thus formed after burrowing but before compaction. laminated siltstone–sandstone (facies 3) this facies varies in grain size from coarse silt to fine sand, but grains up to granule size occur set in a muddy matrix. it forms beds up to 1–2 m thick and is commonly interbedded with facies 2. more than 95% of the grains consist of quartz. other components are plagioclase, muscovite and coaly grains. nodules of framboidal pyrite with a diameter of 0.1 mm occur locally. the cement consists mainly of calcite. the facies is parallel laminated, but structureless intervals are also observed and primary structures are commonly obliterated by bioturbation. laminae are normally 1–2 mm thick, and may be graded from fine sand to coarse silt. thin laminae of shell hash occur locally. the facies is strongly bioturbated. recognizable trace fossils include rare skolithos isp., chondrites isp., rhizocorallium isp. with protrusive spreiten, and ?pygospioides isp. which is very similar to chondrites, but more closely resembles pygospioides isp. as described from the hettangian of niedersachsen by häntzschel & reineck (1968). ?planolites isp. traces occur throughout the facies. body fossils are described under facies 2 (see above). the coarse siltstones and sandstones of this facies probably represent distal offshore storm deposits (pedersen 1985), but the pervasive bioturbation precludes an unequivocal interpretation. variegated mudstone (facies 4) this facies and facies 5 characterise the pankarp member in the upper part of the section (fig. 3). it differs from facies 2 in the red-brown colour and a finer grain size dominated by clay and silt. some levels are light-grey with a greenish tinge. the colour difference from facies 2 is associated with a greater content of iron; the red variety is richer in ferric and poorer in ferrous compounds than the greenish variety. the facies shows some lamination and upwards in the succession thin lenticular silt ripples start to appear. a few indeterminate bivalves and some pyrite-impregnated nodosariid foraminifers have been found. the facies was deposited under low energy conditions, probably in a marginal marine environment as inferred from the impoverished fauna and the scarcity of bioturbation compared to facies 2 and 3. calcareous siltstone and mudstone (facies 5) the facies consists of intimately interbedded, soft, variegated claystone and siltstone of the same type as facies 4, and harder light grey coarse siltstone and very fine sandstone. up to 60–70% of the sediment consists of calcite while quartz, clay and some muscovite constitute the remaining part. the calcite occurs as recrystallized cement and grains of uncertain origin. comminuted coaly fragments occur throughout. the facies shows an almost varve-like grading with 5–17 mm thick beds. the graded beds pass from light grey, calcite-rich clay into red clay, poor in calcite. the lower boundaries of the graded beds are sharp and the tops are flat or gently undulating. some of the thicker, coarser-grained beds show parallel lamination passing into low-amplitude hummocky cross-stratification. these beds also display load structures, groove casts, flatlying folds and wrinkle marks on their upper surfaces. the wrinkle marks are very similar to the kinneya ripples of reineck & singh (1980). trace fossils are scarce, typically represented by scattered 3–4 mm wide subhorizontal burrows, whereas teichichnus isp. and subhorizontal rhizocorallium isp. with protrusive spreiten are found in the upper coarsergrained part of the succession. body fossils are only represented by scattered shell fragments. the fine grain size and the scarcity of trace and body fossils suggest deposition under very low energy conditions in a marginal or high stress marine environment. the graded beds probably represent deposition from storm-induced suspension clouds (pedersen 1985). the upwards increase in grain size and in the frequency of beds displaying parallel lamination and hummocky cross-stratification indicate increasingly storm-influenced 548 deposition. the sum of characters thus indicates deposition under low energy conditions interrupted by sudden influxes of storm-derived sediments. the environment may have been distant offshore or more likely a relatively protected shallow marine area where the available grain sizes were very fine and where the effects of storm events were relatively subtle. the red coloration and the presence of kinneya-type wrinkle marks lend some credence to the latter hypothesis. the wrinkle marks may have been caused by a strong wind blowing over a cohesive, fine-grained sediment covered by only a thin veneer of water, possibly with a microbial mat growing on the sea floor under environmentally stressed conditions (reineck & singh 1980). they are thus indicative of near emergent conditions. as indicated on the geological map (fig. 2), a coal seam is situated slightly above the studied section. this also suggests that facies 5 was deposited in a marginal to non-marine environment. depositional environment facies 1–5 form a regular vertical succession with facies 1 at the base and facies 5 at the top, and interbedding only occurs between facies 2 and 3. the section thus includes a basal muddy sandstone, a lower dark grey unit and an upper unit dominated by light grey and redbrown colours. the sedimentary structures and grain sizes do not show any marked changes and the main mechanisms of transport and deposition seem to have been rather uniform. the whole succession is thus considered to represent an association of genetically related facies. the basal part of the association consists of muddy sandstone (facies 1; only 1.2 m exposed). it is followed with a sharp boundary by a unit dominated by dark grey mudstone with clay-ironstone layers (facies 2), 37.7 m thick, with numerous intercalations of thin siltstone and sandstone beds (facies 3). this unit is overlain with a sharp contact by a variegated mudstone unit, 14.9 m thick, (facies 4), which gradually gives way to a succession of calcareous mudstones and siltstones (facies 5); 15.9 m of this last unit was exposed in the 1970s. the succession is interpreted to reflect changes in relative sea level in an area with uniform subsidence and relatively constant sediment influx. thus, facies 1 represents slow deposition in a well-aerated shallow shelf sea. the muddy sand was originally deposited as alternating thin layers of mud and thicker layers of sand which were thoroughly mixed by bioturbation in the offshore transition to lower shoreface zone close to wave-base. the sharp boundary to the overlying mudstones of facies 2 is interpreted as a ravinement surface caused by combined drowning and transgressive erosion when coarser clastic material was trapped in estuaries and other inshore environments. the main part of the succession represented by mudstone with coarser-grained intercalations (facies 2 and 3) was deposited in deeper offshore areas with periodic influxes of silt and sand from storm-generated suspension clouds. most of the silt and sand beds were thoroughly bioturbated and their identity as storm deposits became less obvious. the unit shows an upwards decrease in the density and diversity of body fossils culminating at the almost barren 35 m level (fig. 3). the top 4 m of the unit are again rich in body fossils. this trend is interpreted to have resulted from transgression and increasing water depth associated with a decrease in oxygenation followed by a regression combined with increasing oxygenation at the sea floor. the variegated mudstones of facies 4 and 5 overlie the dark grey mudstone with a sharp contact (38.9 m in fig. 3) and are somewhat difficult to interpret environmentally. they have the finest grain size of the whole succession indicating very low energy conditions during deposition. the red colour and the occurrence of wrinkle marks or kinneya ripples suggest well-oxygenated, very shallow water conditions. the sediments were possibly derived from erosion of finegrained red beds of triassic age exposed in a nearby source area. the upwards increase in storm siltstones and sandstones suggests coastal progradation, whereas the scarcity of body and trace fossils suggests a marginal marine or high stress environment. the generally fine grain size points to deposition in a sheltered, somewhat enclosed area. the variegated mudstones with storm siltstones and sandstones of facies 4 and 5 thus seem to have been rapidly deposited in a very shallow marine, restricted environment under the influence of storms. sequence stratigraphy the döshult member can be divided into six or seven coarsening-upwards units, about 2–8 m thick, with sharp upper boundaries (fig. 3). they are typical examples of distal parasequences (van wagoner et al. 1990). thin fining-upwards units, 10–20 cm thick, are not assigned any sequence stratigraphic significance but are interpreted as bioturbated storm siltstones and sand549 stones. the parasequences stack into a parasequence set which shows a subtle overall fining-upwards trend accompanied by a distinct decrease in density and diversity of body and trace fossils (0–25 m in fig. 3). the trend is interpreted as an overall backstepping stacking pattern which culminates in the poorly fossiliferous interval between 25 m and 35 m in the upper part of the döshult member (fig. 3). a few thin conglomerates rich in body and trace fossils occur between 35 m and 38.9 m. the dark mudstones of the döshult member (facies 2) are overlain by the pankarp member with a sharp conglomeratic boundary at 38.9 m. this unit comprises about 30 m of variegated, and red-brown, almost unfossiliferous mudstones which contain an upwards increasing number of thin storm siltstones and sandstones. the slowly deposited fully marine muddy sandstones at the base of the section are interpreted to belong to the lower part of the transgressive systems tract. they are capped by a sharp erosional drowning or ravinement surface formed by transgressive marine erosion (tse in fig. 3). the overlying dark mudstones form the upper part of the transgressive systems tract (tst in fig. 3). lowstand deposits cannot be recognised and were probably not deposited in the area. the transgressive marine erosion surface corresponds to the lithostratigraphic boundary between units f-ia and f-ib (both f-i member) of the fjerritslev formation but is slightly younger than in most of the danish basin. the age of the erosion surface is close to the semicostatum–turneri chronozone boundary. a distinct maximum flooding surface cannot be identified on the basis of the available data but a maximum flooding zone is interpreted to occur at about 34 m (mfs in fig. 3). the upper part, from 34 m to the top of the section may represent a simple highstand systems tract. the siderite pebble conglomerate at 38.9 m is not easy to interpret in terms of sequence stratigraphy. it occurs at a marked facies change from dark fossiliferous mudstones (facies 2) to variegated and redbrown almost non-fossiliferous mudstones (facies 4, 5). this change seems to represent a significant environmental change associated with a marked seawards shift in facies and it is possible that it represents a distal sequence boundary. if this is the case then the highstand systems tract of the underlying sequence is a maximum of 5 m thick and consists of dark, uniform mudstones at the top of the döshult member (34–38.9 m in fig. 3). this interpretation is tentatively preferred here and the exposed döshult member thus includes a lower transgressive systems tract (tst), a transgressive surface of erosion (tse), a thick upper transgressive systems tract (tst), a maximum flooding surface or zone (mfs) and a thin highstand systems tract (hst) topped by a distal sequence boundary (sb; fig. 3). the overlying pankarp member probably represents poorly differentiated transgressive and highstand systems tracts. the fossiliferous siderite pebble conglomerate at the döshult–pankarp member boundary may be interpreted as reworked hiatus concretions formed when sediment supply to the basin was shut off during maximum flooding (hesselbo & palmer 1992). this interpretation is, however, considered unlikely due to the marked facies change, the seawards shift in facies and the associated inferred major drop in water depth across the boundary. correlation to the contemporaneous sose bugt member (rønne formation) on bornholm in the baltic sea is hampered by the paralic, poorly fossiliferous nature of that unit (surlyk et al. 1995). dating of the gantofta succession is based on ammonites, ostracodes and foraminifera, whereas the sose bugt member is dated on the basis of pollen in the lower part and a few dinoflagellates in the upper part. a major sequence boundary is situated close to the hettangian–sinemurian boundary in the sose bugt section. this correlates well with a sequence boundary at the base of the döshult member in skåne, below the gantofta section (surlyk et al. 1995). two minor sequence boundaries are identified in the sose bugt section in the lower sinemurian and in the middle upper sinemurian, respectively. the lower sequence boundary occurs at a level roughly corresponding to the top of the muddy sandstone (facies 1) at the base of the gantofta section (1.2 m in fig. 3), whereas the upper one may correlate with the interpreted sequence boundary at the sharp break between dark mudstones (facies 2) and variegated mudstones (facies 4) at gantofta (döshult member – pankarp member boundary; 38.9 m in fig. 3). this correlation may corroborate the interpretation of the erosional boundary between the döshult and pankarp members as representing the distal expression of a sequence boundary. it is remarkable that highstand systems tract deposits are almost absent in the sinemurian sose bugt section which mainly consists of transgressive systems tract deposits. this is thought to be typical of the more proximal, basin margin areas (surlyk et al. 1995) and may also account for the thinly-developed highstand deposits at gantofta which occupied an intermediate basinal position between the paralic setting of the sose bugt member and the offshore danish basin. higher parts of the pankarp member are known from boreholes situated close to the gantofta quarry. the red-brown, var550 551 iegated and light grey mudstones of facies 4 and 5 are overlain by 10–17 m of bluish-grey mudstone. this is followed by about 5 m of sand with an allochthonous coal seam, 5–15 cm thick, representing a marked regression (sivhed 1980). the sand is overlain by about 15 m of red-brown or bluish-grey mudstone. the pankarp member spans the upper sinemurian oxynotum and most of the raricostatum chronozones. the nature of the lower boundary of the sand bed is not known and a sequence stratigraphic interpretation cannot be undertaken on the basis of the available data. a relative sea-level curve constructed on the basis of the sedimentary evolution as interpreted here is shown on figure 4. it is compared with the jurassic eustatic sea-level curves of haq et al. (1988) and hesselbo & jenkyns (1998). the overall trends of the curves are remarkably similar, but the exact ages of the main highs and lows differ somewhat. the haq et al. (1988) and hesselbo & jenkyns (1998) curves show the highest degree of similarity although the latter is more detailed and shows more candidate sequence boundaries and maximum flooding surfaces. the two curves show major sequence boundaries in the uppermost hettangian, uppermost lower sinemurian and uppermost sinemurian, and maximum flooding surfaces in the middle lower sinemurian and middle upper sinemurian. the gantofta curve is simpler due to a combination of uniform facies development and lower biostratigraphical resolution. it differs from the curve of hesselbo & jenkyns (1998) in that they place the main sinemurian sequence boundary at the base or immediately below the base of the obtusum chronozone whereas it occurs within this chronozone at gantofta. the mismatch between the gantofta curve and the two other curves may be due to the basin marginal position and the lack of lowstand deposits at gantofta. the low biostratigraphic resolution prevents identification of possible hiatuses in the mudstone-dominated succession. the eustatic signal may thus be overprinted by tectonism in the fennoscandian border zone, by higher sediment input during transgression and condensation and bypass during regression. hallam (1988) did not give any detailed zonal data for his transgressive and regressive events and his curve is thus difficult to compare with the other curves. conclusions until recently, a lower jurassic, sinemurian marine succession, 70 m thick, was exposed at the gantofta locality in north-western skåne, southern sweden. gantofta represents the only place where it has been possible to study exposed strata of the same facies as the deeplyburied contemporaneous fjerritslev formation of the danish basin. the succession comprises the upper part sb sb sb sbsb sb sb sb mfs mfs mfs mfs mfs high low high low high low after haq et al. (1988) after hesselbo & jenkyns (1998) this paper, gantofta 200 195 ma jamesoni chronozones raricostatum oxynotum obtusum turneri semicostatum bucklandi angulata u pp er u pp er lo w er lo w er pliensbachian stages sea-level changes sinemurian hettangian fig. 4. relative sea-level curve constructed for the gantofta succession compared with the jurassic sea-level curves of haq et al. (1988) and hesselbo & jenkyns (1998); the time scale is after gradstein et al. (1994). an ammonite from the basal muddy sandstone at gantofta suggests a mid-semicostatum chronozone age for this level. the level of the obtusum chronozone is well located and the lower part of the succession has a general early sinemurian age. the age of the postobtusum chronozone beds is not wellknown but ostracode data suggest a late sinemurian age (sivhed 1980). mfs, maximum flooding surface; sb, sequence boundary. of the döshult member and the lower part of the pankarp member, both belonging to the sinemurian–aalenian rya formation. the basal 1.2 m of the gantofta section exposed the uppermost levels of the sand-dominated lower sinemurian part of the lower döshult member. this overall transgressive, fluvial and lacustrine to shallow marine succession is known from temporary exposures at nearby örby where it is 32 m thick (erlström et al. 1999). five genetically related facies are recognised. the lower half of the gantofta section, representing the upper döshult member, is composed of three facies (1–3). the lowermost 1.2 m consists of lower sinemurian bioturbated, richly fossiliferous muddy sandstones (facies 1), interpreted as having been deposited relatively slowly in an offshore to transition zone environment. they are followed with a sharp contact by lower – lower upper sinemurian, dark grey, bioturbated, fossiliferous mudstones (facies 2) with intercalations of siltstones and sandstones (facies 3). the mudstones represent slow, fair-weather deposition below wave base under offshore shelf conditions interrupted by deposition of thin silts and sands from storm-generated suspension clouds. the general low sedimentation rate and the distal, thin nature of the storm deposits is reflected by the pervasive bioturbation and mixing of both facies. the succeeding upper sinemurian succession, referred to the lower pankarp member, comprises marginal marine, variegated mudstones and red-brown calcareous mudstones (facies 4, 5) with an upwards increasing number of storm siltstones and sandstones reflecting general shallowing and progradation of the coastline associated with restriction of the marine circulation. a fossiliferous siderite pebble comglomerate occurs at the boundary between the döshult and pankarp members. the succession encompasses the greater part of two, relatively simple depositional sequences. the basal muddy sandstone is interpreted as belonging to the lower transgressive systems tract of the lower sequence. it is topped by a ravinement surface formed by transgressive marine erosion overlain by a backstepping parasequence set representing the upper transgressive systems tract. a maximum flooding zone is identified close to the top of the dark döshult member mudstones. it is overlain by a thinly developed highstand systems tract topped by an erosion surface marked by the siderite pebble conglomerate. the erosion surface is tentatively interpreted as a distal sequence boundary and the overlying variegated and red-brown pankarp member mudstones belong to the poorly differentiated transgressive and highstand systems tracts of the second sequence. comparison with the sequence stratigraphy of the contemporaneous sose bugt member (rønne formation) of bornholm lends some credence to this interpretation. the gantofta succession records an early sinemurian sea-level rise, a mid-sinemurian highstand, an early late sinemurian sea-level fall followed by a late sinemurian minor rise and subsequent major fall. data from nearby boreholes indicate an end sinemurian – early pliensbachian major rise. the sea-level curve constructed on the basis of the gantofta section is compared with the sea-level curves of haq et al. (1988) and hesselbo & jenkyns (1998) in figure 4. it is remarkable that the three curves show similar overall trends but the sequence boundaries and maximum flooding surfaces are delayed in the gantofta curve compared to the two other curves. this may reflect the basin marginal position of the gantofta section in the fennoscandian border zone with higher sedimentation rates during sea-level rise and condensation or bypass during fall. furthermore, the gantofta curve is much simpler than the hesselbo & jenkyns (1998) curve. this is probably a direct result of the poor biostratigraphic resolution of the gantofta section and the implicit difficulty in identifying hiatuses in the mudstone-dominated succession. acknowledgements we thank ulf sivhed for useful comments, lars b. clemmensen for critically reading an early manuscript version and referees stephen p. hesselbo and gunver k. pedersen for constructive criticism. the study was supported by the carlsberg foundation and the danish natural science research council. references ahlberg, a., sivhed, u. & erlström, m. 2003: the jurassic of skåne, southern sweden. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 527–541 (this volume). bölau, e. 1959: der südwestund südostrand des baltischen schildes (schonen und ostbaltikum). geologiska föreningens i stockholm förhandlingar 81, 167–230. bölau, e. 1973: pankarpslager (lias beta). stratigrafisk definition och facies. geologiska föreningens i stockholm förhandlingar 95, 268–272. erlström, m., sivhed, u. & surlyk, f. 1999: a backstepping fluviatile–paralic–marine succession, sinemurian, lower jurassic, 552 553 skåne, southern sweden. bulletin of the geological society of denmark 46, 1–12. frandsen, n. 1977: aflejringsmiljøer i skånes rhæt–lias, 95 pp. unpublished cand. scient. thesis, københavns universitet, danmark. gradstein, f.m., agterberg, f.p., ogg, j.g., hardenbol, j., van veen, p., thierry, j. & huang, z. 1994: a mesozoic time scale. journal of geophysical research 99, 24051–24074. gravesen, p., rolle, f. & surlyk, f. 1982: lithostratigraphy and sedimentary evolution of the triassic, jurassic and lower cretaceous of bornholm, denmark. danmarks geologiske undersøgelse serie b 7, 51 pp. hallam, a. 1988: a reevaluation of jurassic eustasy in the light of new data and the revised exxon curve. in: wilgus, c.k. et al. (eds): sea-level changes – an integrated approach. society of economic paleontologists and mineralogists special publication 42, 261–273. häntzschel, w. & reineck, h.e. 1968: fazies-untersuchungen im hettangium von helmstedt (niedersachsen). mitteilungen aus dem geologischen staatsinstitut in hamburg 37, 5–39. haq, b.u., hardenbol, j. & vail, p.r. 1988: mesozoic and cenozoic chronostratigraphy and cycles of sea-level change. in: wilgus, c.k. et al. (eds): sea-level changes – an integrated approach. society of economic paleontologists and mineralogists special publication 42, 71–108. hesselbo, s.p. & jenkyns, h.c. 1998: british lower jurassic sequence stratigraphy. in: de graciansky, p.-c. et al. (eds): mesozoic and cenozoic sequence stratigraphy of european basins. sepm (society for sedimentary geology) special publication 60, 561–581. hesselbo, s.p. & palmer, t.j. 1992: reworked early diagenetic concretions and the bioerosional origin of a regional discontinuity within british jurassic marine mudstones. sedimentology 39, 1045–1065. lund, j.j. 1977: rhaetic to lower liassic palynology of the onshore south-eastern north sea basin. danmarks geologiske undersøgelse ii. række 109, 129 pp. michelsen, o. 1975: lower jurassic biostratigraphy and ostracods of the danish embayment. danmarks geologiske undersøgelse ii. række 104, 287 pp. michelsen, o. 1978: stratigraphy and distribution of jurassic deposits of the norwegian–danish basin. danmarks geologiske undersøgelse serie b 2, 28 pp. michelsen, o. 1989: log-sequence analysis and environmental aspects of the lower jurassic fjerritslev formation in the danish subbasin. danmarks geologiske undersøgelse serie a 25, 23 pp. michelsen, o., nielsen, l.h., johannessen, p.n., andsbjerg, j. & surlyk, f. 2003: jurassic lithostratigraphy and stratigraphic development onshore and offshore denmark. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 147–216 (this volume). nielsen, l.h. 2003: late triassic – jurassic development of the danish basin and the fennoscandian border zone, southern scandinavia. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 459–526 (this volume). norling, e. 1972: jurassic stratigraphy and foraminifera of western scania, southern sweden. sveriges geologiska undersökning serie ca 47, 120 pp. norling, e. 1981: upper jurassic and lower cretaceous geology of sweden. geologiska föreningens i stockholm förhandlingar 103, 253–269. norling, e. & bergström, j. 1987: mesozoic and cenozoic tectonic evolution of scania, southern sweden. in: ziegler, p.a. (ed.): compressional intra-plate deformations in the alpine foreland. tectonophysics 137, 7–19. norling, e., ahlberg, a., erlström, m. & sivhed, u. 1983: guide to the upper triassic and jurassic geology of sweden. sveriges geologiska undersökning serie ca 82, 71 pp. pedersen, g.k. 1985: thin, fine-grained storm layers in a muddy shelf sequence: an example from the lower jurassic in the stenlille 1 well, denmark. journal of the geological society (london) 142, 357–374. pedersen, g.k. 1986: changes in the bivalve assemblage of an early jurassic mudstone sequence (the fjerritslev formation in the gassum 1 well, denmark). palaeogeography, palaeoclimatology, palaeoecology 53, 139–168. pieńkowski, g. 1991a: liassic sedimentation in scania, southern sweden: hettangian–sinemurian of the helsingborg area. facies 24, 39–86. pieńkowski, g. 1991b: eustatically-controlled sedimentation in the hettangian–sinemurian (early jurassic) of poland and sweden. sedimentology 38, 503–518. reineck, h.e. & singh, i.b. 1980: depositional sedimentary environments, 549 pp. berlin, heidelberg, new york: springer verlag. reyment, r.a. 1969a: upper sinemurian (lias) at gantofta, skåne. geologiska föreningens i stockholm förhandlingar 91, 208–216. reyment, r.a. 1969b: a note on promicroceras. geologiska föreningens i stockholm förhandlingar 91, 440–442. rolle, f., koch, j.-o., frandsen, n. & surlyk, f. 1979: jurassic environments in the fenno-scandian border zone. symposium on ‘sedimentation jurassique w. européen’. association sedimentologie francais publication speciale 1, 15–31. sellwood, b.w. 1971: the genesis of some sideritic beds in the yorkshire lias. journal of sedimentary petrology 41, 854–858. sellwood, b.w. 1972: tidal-flat sedimentation in the lower jurassic of bornholm, denmark. palaeogeography, palaeoclimatology, palaeoecology 11, 93–106. sivhed, u. 1977: a lower jurassic ostracode fauna in the gantofta brick pit, skåne, southern sweden. sveriges geologiska undersökning serie c 730, 31 pp. sivhed, u. 1980: lower jurassic ostracodes and stratigraphy of western skåne, southern sweden. sveriges geologiska undersökning serie ca 50, 84 pp. sivhed, u. 1981: stratigraphy of the gantofta–katslösa area in scania, sweden. geologiska föreningens i stockholm förhandlingar 103, 249–252. sivhed, u. 1984: lithoand biostratigraphy of the upper triassic – middle jurassic in scania, southern sweden. sveriges geologiska undersökning serie c 806, 31 pp. surlyk, f. & noe-nygaard, n. 1986: hummocky cross-stratification from the lower jurassic hasle formation of bornholm, denmark. sedimentary geology 46, 259–273. surlyk, f., arndorff, l., hamann, n.-e., hamberg, l., johannessen, p.n., koppelhus, e.b., nielsen, l.h., noe-nygaard, n., pedersen, g.k. & petersen, h.i. 1995: high-resolution sequence stratigraphy of a hettangian–sinemurian paralic succession, bornholm, denmark. sedimentology 42, 323–354. troedsson, g. 1951: on the höganäs series of sweden (rhaeto– lias). lunds universitets årsskrift ny följd 2 47(1), 269 pp. van wagoner, j.c., mitchum, r.m., campion, k.m. & rahmanian, v.d. 1990: siliciclastic sequence stratigraphy in well logs, cores, and outcrops: concepts for high-resolution correlation of time and facies. american association of petroleum geologists methods in exploration series 7, 55 pp. 554 manuscript received 23 may 1997; revision accepted 13 october 1999. geological survey of denmark and greenland bulletin 41, 2018, 17-20 17 in organic-rich shales, pores form during oil and gas genesis within organic matter (om) domains. the porosity thus differs markedly from that of conventional reservoir lithologies. here we present the first description of shale fabric and pore types in the lower palaeozoic shales on bornholm, denmark. the pores have been studied using the focused ion beam scanning electron microscope (fib-sem) technique, which allows for high resolution sem images of ion polished surfaces. shale porosity is influenced by many factors including depositional fabric, mineralogical composition, diagenesis and oil and gas generation (schieber 2013). here we discuss some of these factors based on a study of lower palaeozoic shale samples from the billegrav-2 borehole on bornholm (fig. 1) undertaken by henningsen & jensen (2017). the shales are dry gas-mature (2.3% graptolite reflectance; petersen et al. 2013) and have been extensively used as analogies for the deeply buried palaeozoic shales elsewhere in denmark (schovsbo et al. 2011; gautier et al. 2014). the danish lower palaeozoic shale gas play was tested by the vendsyssel-1 well drilled in northern jylland in 2015. gas was discovered within a c. 70 m thick gas-mature, organicrich succession (ferrand et al. 2016). however, the licence was subsequently relinquished, due to a too low gas content. the present study confirms a close similarity of pore development between the shales on bornholm and in the vendsyssel-1 indicating a high porosity within this stratigraphic level throughout the subsurface of denmark. however, the rather different development of porosity in the different shale units presents a hitherto neglected aspect of the palaeozoic gas play in denmark. methods ten samples were selected for thin section and nanoscopic pore analyses based on a screening of 30 samples from the billegrav-2 borehole (fig. 2). total organic carbon (toc) was determined by measuring co2 evolved from the shale fabric and organic nanoporosity in lower palaeozoic shales, bornholm, denmark lucy malou henningsen, christian høimann jensen, niels hemmingsen schovsbo, arne thorshøj nielsen and gunver krarup pedersen denmark 50 km bornholm skåne sweden germany kattegat norwegian–danish basinringkøbing–fyn high a billegrav-2 vendsyssel-1 lower palaeozoic strata caledonian fr well ont jylland fig. 1. distribution of lower palaeozoic strata and wells mentioned in the text. modified from schovsbo et al. (2011). fig. 2. stratigraphy of the billegrav-2 core and overview of samples and fabric types. modified from schovsbo et al. (2011). facies associations in the interval 35–125 m are adopted from the billegrav-1 well described by pedersen (1989); above this level the association is based on the present text. as fm: alum shale formation. dicel: dicellograptus shale. k: komstad limestone. lenticular: lenticular clast-rich mudstone. lin: lindegård mudstone. lithostratigr: lithostratigraphy. l: læså formation. mudsh: mudshale. fabric types: see text. sandstone limestone grey shale m ud st on e le nt icu lar bi om ot tle d sil t-r ich black shale thin section; observed fabric toc / porosity sample pe rio d as so cia tio n m ud sh . as f m o rd ov ici an sil ur ian c am br ian depth (m) fabric k l li th os tr at igr . li n ra st rit es sh ale d ice l sil ts ha le m ud st on e 0 25 50 75 100 125 li th ol og y © 2018 geus. geological survey of denmark and greenland bulletin 41, 17–20. open access: www.geus.dk/bulletin http://www.geus.dk/bulletin 1818 combustion of acid pre-treated samples at 1300°c. porosity was measured in a double-chambered helium porosimeter at the geological survey of denmark and greenland. thin sections with a thickness of about 20 µm were prepared by pelcon material & testing aps. the sem imaging of nano to microscale porosity was performed on cross-sections that were milled and surface polished using a focused ion beam (fib) at the technical university of denmark. in order to minimise erosion of the ion-cut surface, the selected crosssection site was protected with a 3 µm thick layer of platinium. no coating of the imaged surfaces was applied. results each of the stratigraphical units shows a statistically significant correlation between toc and porosity (fig. 3). the rastrites shale at 30–62 m in the borehole is the most porous shale and is characterised by the highest ratio between toc and porosity, whereas the alum shale is the least porous shale, characterised by the lowest ratio between toc and porosity (fig. 3). the dicellograptus shale plots between these trends together with samples from the upper 30 m of the rastrites shale (fig. 3). four shale fabrics are distinguished: (1) a dark-coloured mudstone fabric with high concentrations of om and pyrite, (2) a lenticular clast-rich mudstone fabric, (3) a silt-rich mudstone fabric and (4) a bio-mottled mudstone fabric. the dark-coloured mudstone fabric was observed in five samples and it is the dominant fabric in the alum shale. the fabric comprises a clay-dominated mudstone with variable siltcontent that sometimes contains sand-sized authigenic barite (fig. 4a). the dark colour is due to high contents of dispersed om and pyrite. this fabric is attributed to a generally slow settling of particles in a low-energy depositional environment. the lenticular, clast-rich mudstone fabric is seen in four samples from the alum and rastrites shales (fig. 2). the content of om and pyrite is highest in the dark grey samples and lowest in the pale grey samples (fig. 4a). the typical lenticular clasts range in size from 500 µm to more than 2 mm and are composed of clay and silt-sized material. on a macroscopic scale, the lenticular clasts create a laminated appearance to the shale. the clasts are interpreted as deposited during episodic increases in energy in an otherwise low-energy environment. the silt-rich mudstone fabric (fig. 4b) shows varying concentrations of disseminated silt grains and is observed in four samples from the rastrites shale (fig. 2). more dense accumulations of silt grains in laminae and streaks are typically carbonate cemented. the fabric is assumed to be connected to episodic higher-energy currents in the otherwise low-energy depositional environment. the bio-mottled mudstone fabric (fig. 4c) occurs in three samples from the dicellograptus and rastrites shales. the fabric contains low amounts of om that also tends to be irregularly distributed, both across and along the bedding planes. the distribution reflects the activities of depositfeeding organisms. the fabric is interpreted as deposited in a more oxic marine environment characterised by low om levels and presence of infaunal organisms. pores related to the om vary from simple isolated pores to large pore populations with internally complex structures. isolated pores are usually discrete and equant in shape and can occur both widely disseminated and in more dense populations (figs 4d, e). they also occur in om occupying the space between individual pyrite crystals in framboids (figs 4f, g). the pore-size is usually <100 nm. more dense populations of <50 nm-sized, foam-like pores are also observed. this pore type seems to populate entire om domains, but may also be surrounded by non-porous zones in presumably coherent om domains. a third pore type consists of highly irregular pores with complex internal sub-parts (fig. 4h). this type has a stalactite-like texture with irregular and serrated internal pore surfaces and may have internal fibrous textures resembling wood wool. pores related to the inorganic particles are mainly associated with irregularly shaped grains of quartz and pyrite (fig. 4i). these pores usually appear as discontinuous slits along parts of the grain surface or as curved embayments into the fig. 3. total organic carbon (toc) content versus porosity. arrows represent positive correlation trends (significant at a calculated probability of 0.1) within the stratigraphical units. points in brackets represent data omitted in the correlations. data from vendsyssel-1 are wire-line, log-derived average values (ferrand et al. 2016). po ro sit y (v ol . % ) 0 0 2 4 6 8 alum shale dicellograptus rastrites 3–30 m rastrites 30–62 m billegrav-2: rastrites shale lindegård fm dicellograptus shale alum shale fm vendsyssel-1 (average values): dicellograptus to lower rastrites shale? alum shale formation ( ) toc (wt.%) 12108642 y = 4.5 + 1.8x; r2 = 0.3 y = 4.1 + 0.48x; r2 = 0.7 y = 2.7 + 0.61x; r2 = 0.7 y = 1.7 + 0.26x; r2 = 0.3( ) 19 grain, and they are up to several 100 nm long. other pore types primarily related to inorganic particles are dissolution pores that occur where matrix minerals have become partly dissolved (fig. 4i). this pore type occurs only along the edges of carbonate minerals, and the pores tend to be elongated and irregularly shaped. discussion the variable correlation between the toc content and porosity for the alum, dicellograptus and rastrites shales indicates that pores in both organic and inorganic matter contribute to the total porosity. within each shale unit the toc content correlates with porosity suggesting that pores hosted in organic matter are dominant in all units but with additional contributions from inorganic porosity. a higher contribution of inorganic interparticle pores is seen in the dicellograptus and rastrites shales that add to the overall more porous nature of these shales (fig. 3). the dicellograptus and rastrites shales belong to the mudstone and siltstone associations of pedersen (1989) whereas the alum shale belongs to the mudshale association (fig. 2) and apparently the lithofacies was the main controlling factor of the porosity development. sem images show that the porosity predominantly occurs within amorphous om domains intermingled with the inorganic matrix minerals, rather than as inter-particle pores between the matrix minerals. however, not all om domains contain pores and those that do exhibit considerable variation in quantity, distribution and size of pores. the presence of om in the interparticle spaces cannot be explained entirely by the processes of admixing and subsequent compactional deformation of organic and inorganic fig. 4. micrographs of different mudstone fabrics and pores recognised in the palaeozoic shales. a: dark coloured mudstone fabric intercalated with laminae of lenticular clast-rich mudstone fabric, 119.78–119.80 m (alum shale formation). b: silt-rich mudstone fabric with normal grading, 41.15–41.17 m (rastrites shale). c: biomottled mudstone fabric, presumably chondrites, 86.68–86.70 m (dicellograptus shale). organic pores: d: rounded pores, 74.77–75.01 m (dicellograptus shale). e: sub-rounded pores, 24.78–24.80 m (rastrites shale). f: subrounded to rounded pores in pyrite, 115.63–115.65 m (alum shale formation). g: subrounded to rounded pores in pyrite, 115.63–115.65 m (alum shale formation). h: irregularly shaped and complex pores, 24.78–24.80 m (rastrites shale). inorganic pores. i: irregularly shaped pores surrounding silt to clay-sized grains, 115.63–115.65 m (alum shale formation).  1 mm 1 mm 1 µm a c d g e h f i b 1 mm 500 nm 500 nm 500 nm 500 nm 500 nm 2020 particles (cf. kennedy et al. 2002). instead, it appears that secondary om migrated into interparticle spaces during maturation. this interpretation is supported by observations of well-connected viscous-like om domains, which fill the spaces between matrix minerals. the dominant clay mineral in all the samples is illite (cf. pedersen 1989), which was either a detrital mineral or formed after diagenetic transformation of smectite during burial maturation. it may be assumed that an early migration of secondary om occurred during the temperature interval, which matches the diagenetic transformation of the clay minerals. this relationship between secondary om and diagenetically formed illite was also observed by schieber (2013) in gas-mature samples from the devonian marcellus shale in north america. loucks et al. (2012) suggested that most smectite is transformed to illite during early catagenesis, which supports the observation of presumed migrated om as interparticle fill. comparison with vendsyssel-1 one of the discouraging results of the vendsyssel-1 well was the low porosity and the unfavourable pore distribution in the shales (ferrand et al. 2016). the average toc content and porosity in the vendsyssel-1 well are within the same range as those measured in the billegrav-2 core (fig. 3). sem images of the alum shale from the vendsyssel-1 well show both non-porous om in the mudstone fabric and porous om of presumed secondary origin intermingled with clay minerals (ferrand et al. 2016) similar to the observations from the billegrav-2 core. the similarity suggests that the lower palaeozoic shales known from bornholm are valid analogues for the deeply buried palaeozoic shales in denmark. however, the rather different porosity development in the individual shale units presents a hitherto neglected aspect of the palaeozoic gas play in denmark. conclusions the study shows that the porosities of the lower palaeozoic shales are related to both organic and inorganic matter. the dominating porosity types in all stratigraphical units are those observed within organic matter. a clear relationship between shale fabric and organic nanoporosity has been observed in the lower palaeozoic shales and this indicates that shale composition, depositional environment, and diagenesis have all influenced the porosity development. the toc : porosity relationships in the vendsyssel-1 well are nearly identical to those observed in shales from bornholm indicating a high porosity. the alum shale is a low porous but toc-rich shale whereas the two other shale units studied are low in toc but relatively porous. this observation adds another variable factor to the danish shale gas play (cf. gautier et al. 2014). acknowledgements louise belmonte, formerly at the technical university of denmark, is thanked for providing access to the fib-sem. this paper is a contribution to the geocenter denmark projects 5–2015 and 3–2017. references ferrand, j., demars, c. & allache, f. 2016: denmark – l1/10 licence relinquishment recommendations report. total e&p, memo 1–9 available from: http://www.ft.dk/samling/20151/almdel/efk/bilag/353/1651289.pdf. verified 17.01.2018. gautier, d.l., schovsbo, n.h. & nielsen, a.t. 2014: resource potential of the alum shale in denmark. unconventional resources technology conference (urtec), 25–27 august 2014, denver colorado. spe2014-1931754-ms. 10 pp. henningsen, l.m. & jensen, c.h. 2017: a petrographic analysis of pores and their distribution in palaeozoic organic-rich shale (the alum shale formation, the dicellograptus shale, and the rastrites shale) from the billegrav-2 core, bornholm, denmark,110 pp. unpublished master thesis, university of copenhagen. geus report files 34178 and 34179). kennedy, m.j., pevear, d.r. & hill, r.j. 2002: mineral surface control of organic carbon in black shale. science 295, 657–660. loucks, r.g., reed, r.m., ruppel, s.c. & hammes, u. 2012: spectrum of pore types and networks in mudrocks and a descriptive classification for matrix-related mudrock pores. aapg bulletin 96, 1071–1098. pedersen, g.k. 1989: the sedimentology of lower palaeozoic black shales from the shallow wells skelbro-1 and billegrav-1, bornholm, denmark. bulletin of the geological society of denmark 37, 151–173. petersen, h.i., schovsbo, n.h. & nielsen, a.t. 2013: reflectance measurements of zooclasts and solid bitumen in lower palaeozoic shales, southern scandinavia: correlation to vitrinite reflectance. international journal of coal petrology 114, 1–18. schieber, j. 2013: sem observations on ion-milled samples of devonian black shales from indiana and new york: the petrographic context of multiple pore types. aapg memoir 102, 153–171. schovsbo, n.h., nielsen, a.t., klitten, k., mathiesen, a. & rasmussen, p. 2011: shale gas investigations in denmark: lower palaeozoic shales on bornholm. geological survey of denmark and greenland bulletin 23, 9–12. authors’ addresses l.m.h., energinet, tonne kjærsvej 65, dk-7000 fredericia, denmark; email: lucymalou@gmail.com. c.h.j., region sjælland, alleen 15, dk-4180 sorø, denmark. n.h.s. & g.k.p., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. a.t.n., department of geosciences and natural resource management, university of copenhagen. øster voldgade 10, dk-1350 copenhagen k, denmark. http://www.ft.dk/samling/20151/almdel/efk/bilag/353/1651289.pdf http://www.ft.dk/samling/20151/almdel/efk/bilag/353/1651289.pdf mailto:lucymalou@gmail.com geological survey of denmark and greenland bulletin 31, 2014, 43-46 43 arctic plant remains of weichselian age from the danish north sea ole bennike, jørgen o. leth, jørn bo jensen, niels nørgaard-pedersen and steen lomholt the north sea is a large, shallow epicontinental sea dominated by a sandy bottom that reflects a high-energy environment. little is known about the environmental history of the danish part of this large area during the weichselian, the last ice age. parts of it were glaciated during the last glacial maximum and probably also during older glaciations. shallow parts were dry land, and deeper parts were covered by the sea during ice-free intervals. large, partly ice-dammed lakes also existed. three remains of walrus (odobenus rosmarus) have been radiocarbon dated (möhl 1985; kim aaris-sørensen, unpublished). one of the finds was dated to c. 35 cal. ka bp, whereas the two others finds gave somewhat younger ages of c. 30 cal. ka bp (table 1). knudsen (1985) described the stratigraphy of sediment cores from the western part of the danish north sea. from the roar 41 core, she reported on eemian deposits overlain by weichselian marine deposits with an arctic or boreoarctic foraminiferal fauna. an age of 30–50 cal. ka bp was suggested for this deposit by larsen et al. (2009, fig. 8h). leth (1998) obtained three radiocarbon ages of 43–45 cal. ka bp on marine bivalve shells from two vibrocores from jyske rev. in connection with mapping of sand and gravel deposits in 2012, a number of 6 m long vibrocores were collected by the geological survey of denmark and greenland (geus). during description of the sediment cores, remains of plants were noted; most of them were early holocene. however, a few samples from two sediment cores contained macrofossils of arctic plants and two samples were dated. the aim of this paper is to report on these ages and their implications. the locations of the coring sites are shown in fig. 1. material and methods sediment coring was carried out with a vibrocorer with a 6 m long, 10 cm wide steel core barrel with a pvc tube mounted. coring positions were selected from high-resolution, shallow seismic profiles. the cores were collected in pvc tubes and cut into 1 m long sections and shipped to geus, where they were split and described. a few 1–2 kg samples with plant remains were wet sieved and the residue left on the sieves was analysed using a dissecting microscope. two samples of terrestrial plant remains were dried and dated by radiocarbon accelerator mass spectrometry. a late weichselian deposit lithological logs of the cores are shown in fig. 2. ns 12-213 consists of 100 cm of holocene marine silt and very finegrained sand, underlain by 178 cm indistinctly laminated clay and silt, 180 cm of alternating layers of fine-grained sand, medium-grained sand and silt, 15 cm silt and 71 cm of alternating layers of fine-grained sand, medium-grained sand and silt. the deposits below the marine unit were interpreted as late glacial. plant remains were noted near the bottom of the core. remains of terrestrial plants include leaves of salix herbacea, a leaf of salix phylicifolia, a leaf of betula nana, seeds of empetrum nigrum, a seed of lychnis flos-cuculi, a megaspore of selaginella selaginoides and a stem of distichium sp. © 2014 geus. geological survey of denmark and greenland bulletin 31, 43–46. open access: www.geus.dk/publications/bull 0° 100 km scotland 5°e 0° 5°e 58°n 58°n 56°n 54°n 56°n 54°n denmark norway england the netherlands united kingdom germany 09 13 41 jr >500 200–500 100–200 50–100 40–50 30–40 20–30 10–20 0–10 depth (m) fig. 1. bathymetrical map of the north sea. the dots show the locations of the studied sediment cores (13 = ns 12-2-13, 09 = ns 12-3-09). the green dot shows the location of core roar 41 discussed by knudsen (1985). jr: jyske rev. 4444 (table 2). freshwater plants are represented by potamogeton filiformis, and freshwater invertebrates are represented by cristatella mucedo and fredericella indica. stems of scorpidium scorpioides and drepanocladus s.l. sp. are common; these mosses as well as carex spp. probably grew in wet areas. a sample of salix herbacea leaves gave an age of c. 12.6 cal. ka bp, corresponding to a younger dryas age (table 1). younger dryas deposits are common in denmark, and several submarine deposits from inner danish waters have also been found (e.g. bennike et al. 2004; bennike & jensen ns 12-2-13 56°15.862´ 7°27.819´ lus-10493 salix herbacea 542 12 554–12 746 ns 12-2-09 56°28.478´ 6°27.794´ lus-10494 s. herbacea, d. octopetala 495 33 957–35 010 562026 56°39.3´ 7°39.9´ aar-3291 spisula subtruncata 230 41 136–44 666 562026 56°39.3´ 7°39.9´ aar-3292 macoma balthica 550 42 298–47 250 562028 56°46.3´ 7°33.2´ aar-3293 donax vittatus 310 40 807–45 085 off esbjerg k-3727 odobenus rosmarus 27 033–29 417 table 1. pre-holocene radiocarbon ages from the danish north sea c. 56°30´ c. 7° k-3726 odobenus rosmarus 31 578–36 9401110 ‡ c. 57°12´ c. 8°28´ k-5746 odobenus rosmarus 27 616–33 475a * below core top. † calibrated according to the calib 7.0 program. ‡ core numbers refer to the numbering system used by the marine geologists at geus. § leth (1998). ¶ möhl (1985). a k. aaris-sørensen, unpublished data. § ¶ ¶ § § 1270+– 10 730 ± 75 24 380 ± 620 39 300 ± 1300 41 500 ± 1450 39 200 ± 1100 30 530 ± 300 30 880 26 700 ± 1500 core n. lat. e. long. laboratory species depth age (14c calibrated no. no. b.c.t. (cm)* years bp) age (years bp)† 1 2 3 4 5 cla y sil t vf sand f m d ep th (m ) li th ol og y d ep th (m ) li th ol og y ns 12-2-13 ns 12-3-09 1 2 3 4 5 cla y sil t vf sand f m c vc gravel gr an pe bb table 2. macrofossils of weichselian age from the north sea r: rare, c: common. plants cenococcum geophilum sclerotia 3 15 distichium sp. stems – 1 ditrichum sp. stem – 4 bryum sp. stems – r paludella squarrosa stem – 1 drepanocladus revolvens stems – 2 drepanocladus s.l. sp. stems c r scorpidium scorpioides stems c r polytrichum sp. leaves 1 3 selaginella selaginoides megaspores 1 1 ranunculus sp. achene 1 – betula nana leaf fragment 1 – salix herbacea leaves 7 16 salix phylicifolia leaves 2 – dryas octopetala leaves – 11 empetrum nigrum endocarps 4 – juncus sp. seed – 1 carex spp. achenes 6 14 potamogeton filiformis achenes – 3 animals cladocera indet. shells – r staphylinidae indet. wing 1 – chironimidae indet. head capsules – r trichoptera indet. sclerite – 1 oribatida indet. skeletons – r fredericella indica statoblast 1 – cristatella mucedo statoblast 1 – ns 12-2-13 ns 12-3-09 c. 12.6 ka c. 35 ka fig. 2. lithological logs of core ns 12-2-13 and ns 12-3-09. ns 12-2-13 was sampled at a water depth of 32.0 m and core ns 12-3-09 was sampled at a water depth of 43.5 m. blue: marine holocene. yellow: non-marine pre-holocene. 45 2011; mortensen et al. 2014). however, the present article is the first that reports on a younger dryas deposit from the danish part of the north sea. the deposit shows that the coring site had not yet been transgressed by the sea. a middle weichselian deposit core ns 12-3-09 consists of 75 cm of holocene marine finegrained sand, underlain by 10 cm of holocene marine gravel, 145 cm clay and silt, 238 cm fineand medium-grained sand, 87 cm fine-grained sand with plant remains and 15 cm gravel. remains of terrestrial plants include leaves of salix herbacea, leaves of dryas octopetala, a megaspore of selaginella selaginoides, 2 stems with leaves of ditrichum sp., 1 stem with leaves of distichium sp. and sclerotia of cenococcum geophilum (table 2). wetland species are represented by carex sp., juncus sp., scorpidium scorpioides, drepanocladus revolvens and paludella squarrosa. lake or pond species are represented by potamogeton filiformis, and freshwater invertebrates are represented by cladocerans, chironomid larvae and trichoptera. a sample of s. herbacea and d. octopetala leaves gave an age of c. 35 cal. ka bp, corresponding to a middle weichselian age – prior to the last glacial maximum (table 1). most of the species are known from middle weichselian deposits in denmark, and indicate an open, treeless, tundralike environment with wetlands and lakes. s. herbacea grows in areas with a long-lasting snow cover, whereas d. octopetala prefers areas with little snow cover. the presence of both of these dwarf shrubs indicates that wind-swept hills and places where snow could accumulate were present in the area. non-marine middle weichselian deposits with organic remains are rare in denmark. deposits of broadly the same age as the deposit from the north sea have been discussed by bennike et al. (1994, 2007) and houmark-nielsen et al. (1996); in addition, a number of re-deposited mammal bones and teeth, especially of mammoth (mammuthus primigenius) have been found. the flora and fauna from the other danish sites are similar to those from the north sea. the ålesund interstadial in south-west norway, dated to 28–35 cal. ka bp, is characterised by an arctic vertebrate fauna. pre-holocene bivalve shells and bones as mentioned in the introduction, dating of pre-holocene marine shells from two cores from jyske rev gave ages of 43–45 cal. ka bp (table 1; leth 1998). two of the dated species are boreal and warmth-demanding, which contrasts with indications of arctic conditions during the same time period, as seen in other records from the region. a number 0 50 100 150 200 250 300 350 400 450 500 550 d ep th (c m ) m yti lu s e du lis ce ra sto de rm a ed ul e ar cti ca is la nd ica sp isu la su bt ru nc at a m ac om a ba lth ica d on ax vi tta tu s m ya tr un ca ta ba la nu s c re na tu s ba la nu s b al an us tu rit ell a co m m un is lu na tia a ld er i ac te on to rn at ilis re tu sa tr un ca tu la n uc ul a ni tid os a ch la m ys va ria os tre a ed ul is m ys ell a bi de nt at a ac an th oc ar di um e ch in at um en sis e ns is fa bu lin a fa bu la ab ra a lb a ch am ele a str iat ula d os in ia e xo let a co rb ul a gib ba ec hi no ca rd iu m co rd at um zo ne 1 2 fig. 3. simplified macrofossil diagram of core 562028 from jyske rev (see fig. 1 for location). blue: holocene marine fine-grained sand. yellow: preholocene mediumand coarse-grained sand with pebbles. 4646 of samples from the two cores were analysed in connection with this study. the studied fauna includes a number of boreal species (fig. 3, zone 1). in addition to donax vittatus and spisula subtruncata reported by leth (1998), boreal species are also represented by arctica islandica and cerastoderma edule. these four species are known from holocene and eemian deposits in the region, but not from weichselian deposits. the fauna from the cores also comprises macoma balthica, macoma calcarea, mytilus edulis, hiatella arctica, mya truncata, balanus crenatus and balanus balanus that are boreal and arctic. no species that are confined to the arctic were found, and the fauna can be characterised as of interglacial type. the fauna in zone 2 is typical for holocene fine-grained deposits in the region. there are no reports of interglacial-type deposits from the region from the middle weichselian. for example, in france where the nearest continuous weichselian deposits are found, the middle weichelian was characterised by open vegetation with cold-adapted beetles. hence we suggest that the pre-holocene fauna from jyske rev is of eemian age. that means that the radiocarbon dates should be considered minimum ages. there are numerous examples in the literature of interglacial shells that have yielded non-finite radiocarbon ages which are usually assigned to post-mortem recrystallisation. the deposits with the pre-holocene shells consist of mediumand coarse-grained sand and fine-grained gravel. during periods of low relative sea level, these sandy deposits may have been subject to groundwater flow, which could lead to recrystallisation and introduction of younger carbon. in some areas of the dutch sector of the north sea, bones of cold-adapted mammals are common. the fauna includes terrestrial species such as mammoth (mammuthus primigenius), musk-ox (ovibos moschatus) and reindeer (rangifer tarandus), but also marine species such as walrus, white whale (delphinapterus leucas) and bearded seal (erignatus barbatus). radiocarbon dating of bones from marine mammals yielded middle weichselian ages and non-finite ages (post 2005). we suggest that the finite ages should also be regarded as minimum ages, because we find it doubtful that the sea extended this far south during the middle weichselian, when the global sea level was much lower than at present. we consider it likely that arctic marine mammals lived in the southern north sea during the early weichselian. summary during parts of the weichselian, parts of the danish north sea were land with an open, treeless, tundra-like environment with wetlands and lakes. two radiocarbon-dated arctic floras are dated to the younger dryas and the middle weichselian. a marine fauna from jyske rev gave non-finite middle weichselian ages and we conclude that it is an interglacial-type fauna probably of eemian age. acknowledgement the mapping of sand and graval deposits was funded by the danish nature agency. references bennike, o. & jensen, j.b. 2011: postglacial, relative shore-level changes in lillebælt, denmark. geological survey of denmark and greenland bulletin 23, 37–40. bennike, o., houmark-nielsen, m., böcher, j. & heiberg, e.o. 1994: a multi-disciplinary macrofossil study of middle weichselian sediments at kobbelgård, møn, denmark. palaeogeography, palaeoclimatology, palaeoecology 111, 1–15. bennike, o., jensen, j.b., lemke, w., kuijpers, a. & lomholt, s. 2004: lateand postglacial history of the great belt, denmark. boreas 33, 18–33. bennike, o., houmark-nielsen, m. & wiberg-larsen, p. 2007: a middle weichselian interstadial lake deposit on sejerø, denmark: macrofossil studies and dating. journal of quaternary science 22, 647–651. houmark-nielsen, m., bennike, o. & björck, s. 1996: terrestrial biotas and environmental changes during the late middle weichselian in north jylland, denmark. bulletin of the geological society of denmark 43, 169–176. knudsen, k.l. 1985: foraminiferal stratigraphy of quaternary deposits in the roar, skjord and dan fields, central north sea. boreas 14, 311–324. larsen, n.k., knudsen, k.l., krohn, c.f., kronborg, c., murray, a.s. & nielsen, o.b. 2009: late quaternary ice sheet, lake and sea history of southwest scandinavia – a synthesis. boreas 38, 732–761. leth, j.o. 1998: late quaternary geology and recent sedimentary processes of the jutland bank region, ne north sea, 173 pp. unpublished phd thesis, university of aarhus, denmark. möhl, u. 1985: the walrus, odobenus rosmarus (l.), as a “danish” faunal element during the weichsel ice age. bulletin of the geological society of denmark 34, 83–85. mortensen, m.f., henriksen, p.s. & bennike, o. 2014: living on the good soil: relationships between soils, vegetation and human settlement during the late glacial. vegetation history and archaeobotany 23, 195–205. post, k. 2005: a weichselian marine mammal assemblage from the southern north sea. deinsea 11, 21–27. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: obe@geus.dk geological survey of denmark and greenland bulletin 33, 2015, 53-56 53 follow-up on ujarassiorit mineral hunt finds and outreach activities, south-east greenland majken d. poulsen, holger paulick, diogo rosa, vincent j. van hinsberg, jonas petersen and lærke l. thomsen in connection with fi eld work in south-east greenland in 2014, we took the opportunity to examine the geology associated with potentially valuable mineral occurrences found by local rock collectors. th e initial fi nds were made by local collectors as part of ujarassiorit, which is an annual mineral hunt competition where anyone in greenland can submit samples of rocks they have found i n the countryside for closer examination by the ministry of mineral resources (see www.ujarassiorit.gl). in the tasiilaq region, ujarassiorit resulted in fi nds of corundum, precious metal and base-metal mineral occurrences. our intention was to locate the original sample sites with help from the local rock collectors, describe the geological context and assess the potential for mineral exploration. further work will include laboratory analyses of rock samples and geological reporting. since the base camp in 2014 was located in the settlement kuummiut, we made an eff ort to engage the local people in outreach activities. th e aim was to explain what type of work geologists conduct and how diff erent rocks and minerals can be distinguished. we also visited adjacent communities where the idea was to develop local fi eld guides for teachers, interested citizens and tourists. fig. 1. geological map of the tasiilaq area in south-east greenland, modified from escher (1990) and kolb (2014). © 2015 geus. geological survey of denmark and greenland bulletin 33, 53–56. open access: www.geus.dk/publications/bull ice quaternary isertoq terrane gabbro (c. 1.7 ga) norite, gabbro, diorite, granodiorite (c. 1.85 ga) contact metamorphic anatectic gneiss granite, granodiorite, diorite (c. 1.7 ga) orthogneiss (isertoq) orthogneiss (+opx) amphibolite, ultramafic and metasedimentary rocks isertoq unit (≤1.91 ga) ammassalik intrusive complex marble diorite, tonalite (2.2–1.9 ga) kuummiut terrane brown orthogneiss meta-anorthosite/-leucogabbro ivnartivaq complex (1.955 ga) orthogneiss, migmatitic (c. 2.8–3.0 ga) amphibolite kuummiut unit (<2.2–2.1 ga) helheim unit (>1.9; < 2.2 ga) schweizerland terrane ultramafic rock amphibolite ujarassiorit locality greenland 38°w 66°n 65°30´n 38°w 37°w 36°w se rm ilik 25 km marble belt tonalite belt qiianarteq sermiligaaq schweizerland terrane nordfjord kuummiut aappaluttoq tasiilaq kulusuk tiniteqilaaq nattivit isortoq isertoq terrane ammassalik intrusive complex kuummiut terrane inland ice kangertittivatsiaq 5454 precious metal mineral occurrence north of tasiilaq during fi eld work, the winner of two ujarassiorit prizes, william umerineq from kuummiut, showed the fi eld team the locations of his winning samples. one sample was described by the ujarassiorit jury as a sulphide-bearing, garnet-rich amphibolite rich in platinum-group elements, cobalt and copper; it was collected at ilitti close to nordfj ord at the outer part of kangertittivatsiaq (fig. 1). th e rock contains 44.5% fe, >1% cu, 5280 ppm ni, 2340 ppm co, 804 ppb au and 284 ppm pd. th e other sample was collected near aappaluttoq in the fj ord of sermilik; it had been described as a graphite and garnet gneiss and it had a gold content of 11.1 ppm. at the ilitti locality, a pegmatite lens with coarse-grained calcite was found during fi eld work (fig. 2a). th e origin of the calcite is uncertain; it could come from marble remnants within the host gneiss, from a carbonatite occurrence or it could be of hydrothermal origin. epidote and actinolitebearing, calc-silicate assemblages developed along the contact between calcite and the pegmatite, and small sulphide pods are found locally. such a sulphide pod submitted to the ujarassiorit was confi rmed to be rich in precious metals, cobalt and copper during our follow-up work. at aappaluttoq, the rock collector william umerineq showed us a garnet-rich (garnetite) horizon (fig. 2b). th is horizon seems to belong to one of several discontinuous and foliation-concordant, garnet-rich bodies that are present within yellow-weathering, graphitic mica schist. th ese garnet-rich bodies are probably boudins that formed due to deformation of the garnet-rich rock or its precursor. in some areas, these bodies are closely related to subhorizontal pegmatite sheets, which were also dismembered into boudins. no sulphides were identifi ed in these garnet-rich bodies during fi eld work. however, since umerineq’s sample was very similar in appearance to the garnet-rich bodies that we found, we undertook an extensive sampling eff ort in the area. other samples were collected from silicifi ed domains, veins and veinlets that contain sulphides or gossanised material. laboratory analyses failed to confi rm the presence of a b c d e f fig. 2. a: william umerineq (to the left) at the location where he collected a prize-winning ujarassiorit sample at ilitti in the kangertittivatsiaq fjord area. b: william umerineq standing above the shoreline where he collected another prize-winning, gold-rich ujarassiorit sample near aappaluttoq. the locality is only accessible at low tide. c: william umerineq panning stream sediments to separate possible gold grains on the island of qiianarteq. d: in the isortoq area, dines and rosa jonathansen (first and second from left) took the field team to several of the sites where they had collected rock samples for the ujarassiorit competition. e: the corundum locality was found with help from vittus sakæussen, the 2009 ujarassiorit prize-winner. the geologists are examining ultramafic rocks containing veins with pink corundum; vittus sakæussen is sitting on the ultramafic rock and a light pinkish granitic pegmatite is seen between them. the vein in the ultramafic rock shows several progressive reaction zones. f: biotite, amphibole and pink corundum were formed during the final stages in the development of the metasomatic vein. 55 gold in the garnet-rich rock, but anomalous gold values (up to 463 ppb) were recorded in a few of the veins and veinlets. th is indicates that the original ujarassiorit sample might also have included such a vein or veinlet and that nuggeteff ects could account for the diff erences in gold concentration between that sample and the ones collected during the follow-up in 2014. th e collectors of additional ujarassiorit samples from 1993, 1999 and 2007 with high gold concentrations could not be contacted or accompany us to help locate their sample sites. th e fi rst two samples were erratics whereas the third was in situ, but its precise location is unknown. in order to follow up on these samples, gold panning was carried out in areas with large drainage basins on the island of qiianarteq and on the adjacent peninsula north of qiianarteq (fig. 2c). however, the panning eff ort failed to show any signifi cant gold. follow-up on copper mineral occurrences in the isortoq archipelago th e area around the settlement of isortoq is characterised by an archipelago consisting of banded gneiss with amphibolite bodies hosting granite intrusions. pegmatite dykes and quartz veins are also common. th e fi eld work focused on visiting islands and localities with reported copper mineral occurrences (up to 1.9% cu) in order to describe the geological setting and assess the potential of these occurrences. samples from these localities were submitted to ujarassiorit between 1990 and 2012. some of the sites were visited with the rock collectors, dines and rosa jonathansen (fig. 2d), and this proved to be a good approach to explore several sites in a time-effi cient way. we observed a number of diff erent styles of sulphide occurrences in the area: (1) amphibolite with late-stage brittle fractures with quartz, epidote and pyrite with malachite staining, (2) pyrite within quartz veins, (3) pyrite-chalcopyrite dissemination or veins within pegmatite dykes, (4) sulphide-bearing granitic or gabbroic boudins within banded gneiss, (5) disseminated sulphide in banded gneiss and (6) rounded boulders of sulphide-bearing gabbro in a moraine close to the margin of the inland ice. a soapstone occurrence was also identifi ed that is apparently used by local craft smen as raw material for their work. overall, the fi eld work showed that the known sulphide occurrences are small and commonly lack signs of large-scale hydrothermal alteration. where sulphide occurrences were encountered the visual estimates of pyrite and chalcopyrite concentrations are commonly around 1–2 vol.%. hence, our fi eld work gave little encouragement for additional eff orts to explore for economically signifi cant copper deposits in the area. follow-up on a corundum sample from the ujarassiorit prize-winner in 2009 several occurrences of corundum (al2o3) were investigated, and we focused our eff ort on an occurrence on a small island off immikkeerteq near nattivit (fig. 1). th is corundum locality was visited with the help of the rock collector vittus sakæussen, who initially found it and won the fi rst prize in the 2009 ujarassiorit competition. a preliminary understanding of the formation history of the nattivit occurrence allowed the fi eld team to identify a number of additional corundum occurrences. all occurrences share the same characteristics: corundum occurs where late-stage felsic pegmatites crosscut and interact with metamorphosed ultramafi c rocks (fig. 2e). a progressive metasomatic change of the pegmatite took place at the contact with the ultramafi c body, where progressive subsequently varying mineral zones developed with more and more silica-depleted bulk compositions culminating in a biotite zone, followed by a zone with black amphibole and pink corundum at the centre (fig. 2f). th is sequence of mineral zoning was found as a concentric arrangement along the length of the pegmatite dyke with successively increasing development as the dyke protrudes farther into the metamorphosed ultramafi c rock. th e corundum crystals are generally large (locally up to 5 cm in diameter) and anhedral with minimal fracturing. th is mode of occurrence as isolated, large grains suggests that metasomatic replacement took place under conditions where element mobility was high. th e source of the aluminium appears to be plagioclase in the pegmatite with removal of silica by interaction with the ultramafi c rock which ultimately led to al2o3 saturation. th e process of formation is local and linked to element exchange between ultramafi c rocks and late-stage pegmatite dykes. both these lithologies are common throughout the fi eld area, and metasomatic reaction zones between them are ubiquitous. however, most of the reaction zones only developed biotite, without black amphibole or corundum. older generations of metasomatic interaction between pegmatites and ultramafi c rock units are present, but these lack corundum. th e metamorphosed ultramafi c rocks are mainly sills or dykes that occur as intrusions in a tonaliteamphibolite gneiss basement. black amphibole rims developed at the contact between the intrusion and the gneiss, but no other mineral zones were observed. general outreach and education th e geology near kuummiut, tiniteqilaaq, sermiligaaq, kulusuk and tasiilaq was studied and rock samples collected as 5656 authors’ addresses m.d.p., geological survey of denmark and greenland, nuuk office, kivioq 2, 3900 nuuk, greenland; e-mail: madp@geus.dk h.p. & d.r., geological survey of denmark and greenland, øster voldgade 10, dk-1350, copenhagen k, denmark. v.j.v.h., department of earth & planetary sciences, 3450 university street, montreal, quebec, canada h3a 2a7. j.p.& l.l.t., ministry of mineral resources, postbox 930, imaneq 1a, 201, 3900 nuuk, greenland. a b fig. 3. a: a geologist from geus teaching children in kuummiut how to use a hand lens and how to recognise different minerals. b: robert umerineq enthusiastically studies a rock sample. photographs: jakob lautrup a basis for local geological fi eld guides. th e intention is to engage school teachers and other interested citizens to develop a better understanding of the local geology, the science of geology and geological fi eld work in general. th e local people are already familiar with their local environment and landscapes, and this will allow them to interpret what they see, and teach this to the next generation. th e fi ve communities represent diverse geology and cover a wide spectrum of themes, which taken together, can tell a story of a dynamic geological history of this part of southeast greenland. a variety of geological structures and processes can be studied, such as relations between intrusions at kulusuk and tasiilaq, contact metamorphism at tiniteqilaaq, deformation and metamorphism at kuummiut and an igneous fractionation series at sermiligaaq. th e smaller settlements are well suited for geological outreach because they show the most interesting igneous and metamorphic geology, whereas a fairly uniform geology is exposed around the town of tasiilaq. we collected samples that can provide mineralogical, petrological and geochronological information, as well as data on geochemical compositions and thin section photographs for the fi eld guides. th e fi eld guides will be written following analyses of the rock samples, and will contain background information on geological processes and features. shortly aft er the beginning of the school year in 2014, a geology theme day about basic geology with demonstrations of the typical rock types in kuummiut was held at the primary school in kuummiut for 6–12 years old children (fig. 3). furthermore, a community information meeting was arranged at the school in kuummiut, where we presented the aims of the fi eld activities in the tasiilaq area. th e meeting and presentations were well-received and more than 30 citizens of kuummiut attended. conclusions th e experience from this fi eld season shows that the followup of mineralised samples submitted to the ujarassiorit programme is best carried out with the help of local rock collectors. th is has proved to be the most effi cient way to quickly fi nd the right areas, since the limited time in the fi eld makes it important to get to the right locations as quickly as possible. also, we found that engaging the local community members in the fi eld work contributed to a better understanding of what geologists are doing in the fi eld and how this work may ultimately benefi t the greenlandic community. our experience shows that especially children enjoy exploring their natural surroundings and are eager to learn about rocks and the work carried out by geologists. field guides and further outreach will be appreciated by the local people and will hopefully lead to a self-sustainable situation where school teachers and interested citizens acquire suffi cient geological background knowledge to explain their local geology. acknowledgements th is study is a joint project fi nanced by the government of greenland and the geological survey of denmark and greenland. we are grateful to the rock collectors who shared their local knowledge with us and to the people of kuummiut and isortoq for support and for their enthusiasm for our work. hans kristian olsen kindly commented on the manuscript. references escher, j.c. 1990: geological map of greenland, 1:500 000, skjoldungen, sheet 14. copenhagen: geological survey of greenland. kolb, j. 2014: structure of the palaeoproterozoic nagssugtoqidian orogen, south-east greenland: model for the tectonic evolution. precambrian research 255, 809–822. review article | short rasmussen & dybkjær 2020: geus bulletin 44. 4618. https://doi.org/10.34194/geusb.v44.4618 1 of 6 the lower miocene flint conglomerate, jylland, denmark: a result of the savian tectonic phase erik skovbjerg rasmussen1*, karen dybkjær1 1geological survey of denmark and greenland (geus), copenhagen, denmark abstract the early miocene was an important period for the development of the eastern north sea. tectonism in north-west europe resulted in uplift of the scandinavian mountains, reactivation of salt structures, inversion of old graben structures and deposition of the most coarse-grained deposits in the danish pre-quaternary succession. some of these deposits were later cemented into conglomerates. the deposits are common in the fluvial parts of the billund formation (aquitanian) and the basal transgressive lag of the late aquitanian  – burdigalian klintinghoved formation capping the billund formation. questions remained as to the age of these deposits and what they infer about tectonic events in the region. this study reviews the geology of the flint-dominated conglomerates and presents the first dates for a sample of these unique deposits. we observe grain sizes up to 5 cm diameter. palynological analyses place the sample as early miocene. some samples from the area have suggested a local source near active salt structures, associated with the uplift of the pre-neogene sedimentary successions. we suggest that the common occurrences of flint clasts in the lower miocene succession reveal significant erosion of upper cretaceous and danian chalk, likely associated with the uplift of the scandinavian lowlands during the savian tectonic phase, early miocene. introduction the existence of a conglomerate dominated by flint in central and western jylland, western denmark, has been known for many years. samples of it are commonly caught by rock dredging in the north sea, particularly at jyske rev off the coast of thyborøn (fig. 1b). the conglomerate was described by ødum (1968) as “et groft konglomerat […] af ganske overvejende flintrullesten i en grundmasse af sandsten,” translated as “a coarse-grained conglomerate […] consisting mainly of flint gravel in a matrix of sandstone.” he interpreted the conglomerate as having originated from an unknown ‘palaeogene’ formation located in the skagerrak area. extensive mapping of aquifers in jylland during the past two decades has contributed to a wealth of information about the neogene succession in denmark. detailed studies of outcrops, drilling and analysis of more than 50 boreholes, acquisition of high-resolution seismic data and the development of a solid biostratigraphic framework for the succession now exist (e.g.  dybkjær & rasmussen 2000, 2007; dybkjær & piasecki 2010; rasmussen et al. 2010). *correspondence: esr@geus.dk received: 03 mar 2020 accepted: 21 aug 2020 published: 02 nov 2020 keywords: biostratigraphy, flint conglomerate, north sea, miocene, tectonism geus bulletin is an open access, peerreviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: michael bryld wessel fyhn (geus, denmark) reviewed by: henrik friis (aarhus university, denmark) and one anonymous reviewer funding: see page 6 competing interests: none declared additional files: none provided https://doi.org/10.34194/geusb.v44.4618 mailto:esr@geus.dk rasmussen & dybkjær 2020: geus bulletin 44. 4618. https://doi.org/10.34194/geusb.v44.4618 2 of 6 www.geusbul let in.org these studies revealed a number of conglomerates, most of which, if not all, correlate with gravel layers in the region. the gravel layers formed during the deposition of the lower miocene billund formation and the basal part (transgressive lag) of the succeeding klintinghoved formation. the transgressive lag of the klintinghoved formation, which is commonly cemented, is thought to be the source for these conglomerates found in the quaternary succession in jylland, denmark. in this study, we review the history of these conglomerates and present the first dates for a sample of the unique flint conglomerate first defined by ødum (1968) from offshore jylland, west denmark. we relate these dates to the gravel layers from onshore jylland. in doing so, we begin to unravel the geological significance of this conglomerate, which is a part of the most coarse-grained deposits of the billund and klintinghoved formations in the danish stratigraphic column. geological setting by the end of the early permian, during the final phase of the formation of pangea, the north sea area was characterised by strong earthquake activity and volcanism, and many of the structural elements, including the skagerrak platform, the sorgenfrei–tornquist zone, the norwegian–danish basin and the ringkøbing–fyn high, were formed (ziegler 1990). this period was succeeded by thermal subsidence, which resulted in the formation of a regional depression, the so-called north permian basin (ziegler 1990). the basin was located just north of the palaeo-equator, which formed optimal conditions for an extremely dry climate and a desert environment. during the late permian, this basin was flooded several times and evaporites were deposited. the dry climate continued during the triassic and was characterised by desert environments with the deposition of thick alluvial successions along the southern margin of scandinavia. by the end of the triassic, a humid climate was established, and a regional transgression commenced. marginal marine sedimentation dominated the danish area in the latest triassic and earliest jurassic time, but continued transgression resulted in a fully marine depositional environment in the early jurassic. the north sea basin and scandinavia were strongly affected by mid-jurassic updoming and succeeding rifting activity (e.g. underhill & partington 1993; andsbjerg et al. 2001; andsbjerg & dybkjær 2003). in the aftermath of this tectonically active period, the basin underwent thermal subsidence during most of the cretaceous. consequently, the basin was filled with mudstone and chalk during the cretaceous (surlyk et al. 2013 and references therein). chalk deposition culminated in the late cretaceous, where the influx of siliciclastic sediments to the basin was limited due to a very low relief of the hinterlands and a high global sea level. the relief of parts of southern norway was in the order of c. 100 m (sømme et al. 2019), and most of the scandinavia was covered by marine chalk and mud deposits (jarsve et al. 2014). periodically, however, due to inversion tectonism (alpine orogen), local areas were uplifted at the margin of the scandinavia, and smaller deltas prograded into the north sea basin (erlström 1994). the overall marine depositional setting established in the cretaceous continued during the paleocene. however, the portion of scandinavia facing the north atlantic was exposed to uplift associated with the opening of the north atlantic and a topography of more than 1000 m existed by the end of the paleocene (sømme et al. 2019). uplift resumed in the late eocene, fig. 1 two palaeogeographic reconstructions of the early miocene eastern north sea. approximate locations of the samples used in this study are shown. a: early part of the miocene transgression (c. 22.5 ma). parts of the ringkøbing–fyn high and topographic highs above salt diapirs formed islands in the miocene sea. on the islands formed by salt diapirs, cretaceous and danian chalk was exposed and formed coastal cliffs. b: progradation of the billund delta-system during the early miocene (c. 22 ma). the quartz clasts of the flint conglomerate were transported to the area of west of thyborøn by the western river system that had its outlet in the west of jylland. modified from rasmussen et al. (2010). a b km ringkøbing – fyn high skagerrak ver eksyj thyborøn brøndum salten addit/ voervadsbro km100100 fur hostrup https://doi.org/10.34194/geusb.v44.4618 http://www.geusbulletin.org rasmussen & dybkjær 2020: geus bulletin 44. 4618. https://doi.org/10.34194/geusb.v44.4618 3 of 6 www.geusbul let in.org and prograding wedges from scandinavia formed in the north sea basin in the early oligocene. by the end of the oligocene and in the early miocene, regional scale inversion tectonism occurred in central europe (pharaoh et al. 2010 and references therein). this inversion also influenced the north sea as seen in inversion of the sole pit axis and of the central graben (rasmussen 2009; green et al. 2017). uplift of the southern scandinavia is also indicated by significant sediment influx into the basin (rasmussen 2004) and revealed by exhumation studies (e.g. japsen et al. 2007). consequently, most of the present-day denmark became land during the early miocene (fig. 1; rasmussen et al. 2010). major reorganisation of the tectonic regime during the mid-miocene resulted in flooding of the early miocene deltas in the north sea, and marine conditions were re-established. during the late miocene, delta progradation from scandinavia resumed both in the north sea and in the northern north atlantic. materials and methods during numerous field campaigns in the past 20 years, 25 outcrops have been studied and logged by conventional facies analysis. in addition, a total of 50 boreholes across jylland have been described for sediment texture. this resulted in the establishment of a new lithostratigraphy for the onshore danish miocene succession (fig. 2; rasmussen et al. 2010). a chronostratigraphic framework for the miocene succession was established based on the studies of the dinoflagellate cyst (dinocyst) assemblages in more than 2000 sediment samples from both outcrops and boreholes (dybkjær & piasecki 2010). for this study, we conducted palynological (dinocyst) analyses on a sample of flint conglomerate from jyske rev, offshore jylland, west denmark. the conglomerate contained a sea-urchin of the late cretaceous genus galerites and both angular and rounded flint clasts. this piece is classed as ‘danekræ’ (dk 862; fig. 1b), a danish classification for the preservation of natural historical objects of unique scientific value. in addition, several series of samples have previously been analysed from cemented sandstone from brøndum gravel pit, onshore western denmark and the succession comprising the flint-rich gravel layer from the cliff at hostrup, north-west jylland, and from fluvial, gravel-rich deposits from the addit and voervadsbro pits and the salten section (dybkjær and piasecki 2010 and references therein; fig. 1b). the sediment samples were processed using standard palynological preparation methods, including treatment with hcl to dissolve calcareous matter and with hf to dissolve silicates. the acid treatment was followed by neutralisation with water, brief oxidation with koh and sieving on 20 µm nylon mesh. the organic residues, including the dinocysts, were then mounted on a glass slide using glycerin gel and studied using a normal light microscope. in the following sections, we review the previously published understanding of the age and formation of the miocene conglomerates, supplemented by our new observations and dates from sample dk 862. the miocene conglomerates during the oligocene–miocene transition, a dramatic change occurred in the depositional environment in the eastern north sea basin. from a dominance of pelagic and hemipelagic fine-grained sedimentation that lasted from the paleocene to the oligocene (heilmann-clausen m åd e ri be g ro up m io ce ne lo w e r m id dl e n eo ge ne 15 20 25 cha�an langhian burdigalian aquitanian o lig oc en e u pp er pa la eo ge ne pe rio d sw ne epochma age lithostra�graphy marine silt and clay marine sand brackish water silt and clay fluvial sand and gravel hiatus coal hodde fm arnum fm stauning mb odderup fm bastrup fm fasterholt mb vandel mb resen mb klin�nghoved fm vejle fjord fm brejning fm brejning fm skansebakke mb kolding fjord mb øksenrade mb billund fm hvidbjerg mb not included in this study addit mb fig. 2 lower miocene lithostratigraphy of denmark. note that the billund (including the addit member) and klintinghoved formations form the lowermost part of the lower miocene. modified from rasmussen et al. (2010). https://doi.org/10.34194/geusb.v44.4618 http://www.geusbulletin.org rasmussen & dybkjær 2020: geus bulletin 44. 4618. https://doi.org/10.34194/geusb.v44.4618 4 of 6 www.geusbul let in.org et al. 1985; nielsen et al. 2015), deposition of a vast sand and gravel-rich delta occurred in the early miocene (rasmussen et al. 2010). in this delta, gravel and cobbles were transported into the marine part of the north sea basin. the gravel clasts are composed of quartz, quartzites and flint, some of which originate as far away as jotunheim in present-day norway and areas in central sweden (olivarius et al. 2014). the gravel clasts studied here are up to 5 cm in diameter (figs 3a, b and c). some of the flint cobbles are rounded and show evidence of long transport (fig. 3d), whereas flint clasts found in the north-western part of denmark clearly indicate short transport distances, as indicated by the well-preserved sea urchins (galerites) and angularity of some clasts (figs 3a and b). it cannot be excluded that rounded flint clasts were also formed on beaches fringing local salt structures, but the coexistence of rounded flint clasts, rounded quartz and quartzites of the same grain size indicates that a larger source area, i.e. scandinavia. as such, the wide geographical distribution of the rounded flint clasts points towards a regional source that is not restricted to local salt structures. coarse-grained clasts are commonly found in channel thalweg deposits of the addit member of the billund formation (fig. 2), along erosional boundaries within fluvial systems, or concentrated within a basal conglomerate of the klintinghoved formation that was formed during the early miocene flooding of the billund delta system. all of the palynological samples studied here reveal an early miocene age (fig. 4). the danekræ dk 862 is dated to the early aquitanian, earliest early miocene, dinocyst zone chiropteridium galea of dybkjær and piasecki (2010). the transgressive lag capping the billund delta system is dated to the latest aquitanian – earliest burdigalian, early miocene, dinocyst zone thalassiphora pelagica of dybkjær and piasecki (2010). this places the conglomerate in the lower miocene billund and klintinghoved formations (fig. 2). depositional environment it is likely that the coarse-grained lower miocene deposits were transported in fluvial systems, sourced from scandinavia (present-day southern norway and central sweden). the catchment area covers c. 250 000  km2 (olivarius et al. 2014), and during the earliest early miocene, a predominance of braided river systems characterised the landscape across jylland (rasmussen 2014). the climate was warm and humid (larsson et al. 2011). the river systems were perennial with channels up to 12 m deep and 800 m wide. the eastern portion of the miocene north sea was storm influenced, so the delta systems were wave dominated. the delta systems prograded into a shallow sea, c. 100 m deep (fig. 1). locally, islands were present in northern jylland (fig.  1a). these islands were formed by salt diapirism, similar to the present-day helgoland, german north sea, and exposed paleocene clay, diatomite and cretaceous and danian chalk. large-scale tectonism the abrupt change in the depositional setting, from deep marine clay to deltaic deposits at the oligocene– miocene transition in the eastern north sea, was caused by regional compression associated with alpine tectonism in the savian phase (rasmussen 2004; pharaoh et al. 2010). regional inversion associated with this fig. 3 flint conglomerate (dk 862) and loose clasts found in the lower miocene billund and klintinghoved formations. a: flint conglomerate with a well-preserved sea urchin (galerites) of late cretaceous age, indicated by an arrow. b: angular clast of flint c. 2 cm (arrow). c: elongated flint clast from the fluvial addit member of the billund formation. d: quartzites, quartz, rock fragments and flint found in the basal gravel lag of the klintinghoved formation. note the diameter of the rounded clasts is up to 5 cm and 7 cm. a b c d https://doi.org/10.34194/geusb.v44.4618 http://www.geusbulletin.org rasmussen & dybkjær 2020: geus bulletin 44. 4618. https://doi.org/10.34194/geusb.v44.4618 5 of 6 www.geusbul let in.org phase in central europe resulted in the weald and sole pit anticlines in the uk and up-doming of the igor ridge in the danish north sea (vejbæk et al. 2006; rasmussen 2009; knox et al. 2010; green et al. 2017). the formation of the salt structures, exposing chalk in the early miocene, coincided with this regional inversion tectonism in north-west europe. the sea urchins and angular flint clasts presented here, and those found around salt structures in northern jylland, also place the formation of these structures during the earliest early miocene. therefore, it is likely that parts of the norwegian–danish basin were also inverted during this compressional phase, similar to the igor ridge in the danish central graben and the sole pit and weald basin in the united kingdom. however, due to later erosion, direct evidence was lost. the occurrence of precambrian quartzite clasts, with a diameter of up to 5 cm, reveals also that parts of scandinavia were uplifted during the savian phase. clasts of that size can only be transported in river systems with high flow velocities and a gradient higher than river systems filling a tectonically quiescent basin. although there has not been a quantification of the content of flint in the conglomerates and gravel layers, the common occurrence of flint indicates that a significant portion of the deposits originated from chalk that covered parts of present-day norway and sweden. palaeorelief of southern norway estimated by sømme et al. (2019) reveals a very low topography during the late cretaceous, and given higher sea-level during the cretaceous, most of southern norway was probably submerged at that time (jarsve et al. 2014). a unique conglomerate the conglomerate sample examined here and the associated gravel layers containing both flint, quartz and quartzite are unique in the pre-quaternary succession of denmark. such mixed conglomerates or gravel deposits containing large clasts up to 5 cm in diameter do not occur at any other levels. similarly, large clasts found within the basal part of the hodde formation (koch 1989) consist solely of flint and are probably associated with local salt structures. the presence of quartzite clasts that have their origin in scandinavia indicates long transport distances. this implies a high gradient and discharge of the river systems capable of transporting the clasts. in modern environments, such rivers are found in tectonically active areas. low relief in southern scandinavia during the late cretaceous permitted chalk and clay deposition here (jarsve et al. 2014). some of this chalk was reworked during cenozoic uplift, e.g. during the paleocene and particularly in the early miocene as indicated by the conglomerate described here. we infer that this part of the miocene seems to have experienced significant uplift of the scandinavian mountains and significant erosion of the cretaceous deposits comprising mainly chalk with flint nodules, which, at that time, covered present-day southern norway and central sweden. acknowledgements we would like to thank jacob lind bendtsen and stefan sølberg for help drafting the figures. the lab-technicians anette ryge and charlotte olsen are thanked for the processing of the palynological samples. henrik friis and an anonymous reviewer are thanked for constructive comments on the manuscript. fig. 4 dinocysts from the lower miocene succession onshore denmark (left) and the offshore flint conglomerate (dk 862; right). a: well preserved, whole specimen of homotryblium tenuispinosum from a nearby lower miocene outcrop for comparison. b: h. tenuispinosum fragment from the flint conglomerate (dk 862). c: well preserved, whole specimen of chiropteridium galea from a nearby lower miocene outcrop for comparison. d: c. galea fragment from the flint conglomerate (dk 862). a b c d 20µm 20µm 20µm20µm https://doi.org/10.34194/geusb.v44.4618 http://www.geusbulletin.org rasmussen & dybkjær 2020: geus bulletin 44. 4618. https://doi.org/10.34194/geusb.v44.4618 6 of 6 www.geusbul let in.org additional information funding this study was funded by internal funds at geus. authors’ contributions esr: responsible for sedimentology and regional geology, drafted the main part of the paper. kd: conducted palynological analyses, drafted the palynological methods and text related to dating of the conglomerate. references andsbjerg, j. et al. 2001: divergent development of two neighbouring basins following the jurassic north sea doming event: the danish central graben and the norwegian–danish basin. in: martinsen, o.j. & dreyer, t. 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https://doi.org/10.1002/9781118920435.ch15 https://doi.org/10.1002/9781118920435.ch15 https://doi.org/10.1016/j.palaeo.2011.05.003 https://doi.org/10.1016/j.palaeo.2011.05.003 https://doi.org/10.2110/jsr.2015.40 https://doi.org/10.2110/jsr.2015.40 https://doi.org/10.1016/j.marpetgeo.​2013.09.010 https://doi.org/10.1016/j.marpetgeo.​2013.09.010 https://doi.org/10.1016/j.gloplacha.2003.08.004 https://doi.org/10.1016/j.tecto.2008.10.025 https://doi.org/10.1016/j.tecto.2008.10.025 https://doi.org/10.2110/jsr.2014.24 https://doi.org/10.1111/j.1365-3091.​2004.00681.x https://doi.org/10.1111/j.1365-3091.​2004.00681.x https://doi.org/10.34194/geusb.v22.4733 https://doi.org/10.34194/geusb.v22.4733 https://doi.org/10.3389/feart.2019.00303 https://doi.org/10.1016/j.cretres.2013.08.006 https://doi.org/10.1144/0040337 https://doi.org/10.34194/geusb.v13.4962 the lower miocene flint conglomerate, jylland, denmark: a result of the savian tectonic phase abstract introduction geological setting materials and methods the miocene conglomerates depositional environment a unique conglomerate acknowledgements additional information references figures fig. 1 two palaeogeographic reconstructions of the early miocene eastern north sea. approximate loca fig. 2 lower miocene lithostratigraphy of denmark. note that the billund (including the addit member fig. 3 flint conglomerate (dk 862) and loose clasts found in the lower miocene billund and klintingh fig. 4 dinocysts from the lower miocene succession onshore denmark (left) and the offshore flint con geological survey of denmark and greenland bulletin 31, 2014, 59-62 59 six years of petroleum geological activities in north-east greenland (2008–2013): projects and a view of the future jørgen a. bojesen-koefoed, peter alsen and flemming g. christiansen the deadline for applications to the first licence round for petroleum exploration offshore north-east greenland was 15 december 2012. the round was restricted, allowing only members of the kanumas consortium to be operators (bp, chevron, exxon, jogmeg, shell and statoil). nunaoil is also part of kanumas, but it is a carried, non-operator partner. an ordinary licensing round followed shortly after with a deadline on 15 october 2013. at the end of 2013, pre-round licences were awarded and the results of the ordinary round are expected in 2014. irrespective of the outcome of future exploration activities, this milestone marks an important step in greenland’s long way towards becoming an oil nation. it also offers an opportunity to summarise petroleum-directed activities in northeast greenland since 2008, through which the geological survey of denmark and greenland (geus) has provided crucial knowledge to both the industry and the greenland bureau of minerals and petroleum (bmp) in their preparations for the licence rounds. background almost a decade ago, a licence round offshore north-east greenland, including the danmarkshavn basin (figs 1, 2), was discussed between the kanumas partners and the bmp, and the details of timing and block locations were negotiated over several years. in keeping with this, geus initiated discussions with the petroleum industry to clarify its needs in order to be able to provide optimal service to both companies and authorities when required. concurrently, the united states geological survey’s (usgs) ‘circum arctic resource appraisal’ directed focus to subjects that required further study. the north-east greenland part was carried © 2014 geus. geological survey of denmark and greenland bulletin 31, 59–62. open access: www.geus.dk/publications/bull fig. 1. map of north-east greenland, showing the offshore licence area, onshore field work areas (colour-coded according to year) and the locations of core holes. inland ice danmarkshavn basin store koldewey wollaston forland hold with hope jameson land wandel sea basin 2012 2008 2013 2009 2011 2010 18 15 14 19 17 16 13 11 12 10 9 8 4 2 3 1 7 6 5 kilen tertiary basalts wandel sea basin carboniferous–cretaceous sed. jameson land basin devonian basin palaeozoic shelf sediments palaeozoic trough sediments proterozoic sediments and volcanics caledonian fold belt proterozoic basement 7 core hole block number onshore field work awarded block greenland sea peary land 6060 out in collaboration with geus (christiansen et al. 2006; gautier 2007; gautier et al. 2011; christiansen 2011). the seismic coverage of the north-east greenland shelf is scattered and there are no wells. therefore an understanding of the onshore geology is indispensable to construct analogues for the offshore basins. geologists from danish research institutions have worked in ne greenland for over a century, and the accumulated sample and knowledge base at geus is the largest existing geological database for greenland. in 2008, a project was set up at geus to systematise relevant available data and samples from this database, which contains results from onshore field work, core drilling and subsequent analyses. the aim was to use this material as a starting point for addressing key risks and uncertainties for future offshore exploration. risks for offshore ne greenland petroleum exploration include in broad terms: (1) the distribution, quality and correlation of the main upper jurassic – lower cretaceous source rock units, which have an important bearing on the nature of the petroleum products generated, (2) the nature and stratigraphic distribution of potential reservoir and seal rocks (and possible source rocks), primarily in the rather poorly known cretaceous succession, which is very thick according to the only published seismic interpretation of the danmarkshavn basin (hamann et al. 2005) and (3) the subsidence and exhumation history of north-east greenland. project portfolio in late 2007, a collaboration project, ‘petroleum geological studies, services and data in east and north-east greenland’, was initiated between geus and a number of oil companies. from the start, only little more than a handful of companies chose to participate, but soon the number grew to approximately twenty. the activity was designed with an open end, i.e. to continue for as long as industry interest would last, and it is now entering its seventh year. participation is open to all companies that actively explore for and exploit petroleum. upon entry into the group of partners, each company gets: (1) a gis compilation summarising all available relevant information from north-east greenland, including geological and palaeogeographical maps, data on >17 000 samples, photographs, logs, potential field data, etc. from onshore north-east greenland from 70° to 82°n, (2) three reports on ‘petroleum systems’, ‘uplift’ and ‘seabed features’ mainly based on existing knowledge but including some new and reprocessed data, (3) invitation to semi-annual workshops at geus of which the 12th took place in november 2013, (4) access to data, results, power point presentations, etc. in addition, project partners are invited to participate in a number of projects dedicated to specific topics defined by geus. at the end of 2013, nine such projects, including a field excursion to north-east greenland have been carried out, and an additional four new projects have recently been proposed (table 1). field work and shallow core drilling onshore field work, often including shallow core drilling down to approximately 235 m, has been carried out every summer since 2008 (fig. 1). each year’s field campaign in north-east greenland lasted 3–6 weeks and comprised 20–30 participants. an overview of the 2008 field season, including preliminary results of the drilling of the blokelv-1 core hole was presented by bojesen-koefoed et al. (2009). a principal target for shallow core drilling was the upper jurassic – lower cretaceous source rock succession, which is the local equivalent of the deposits that have generated the main part of the petroleum found in the north sea and its northward extension. the imperfect knowledge of this succession was one of the key problems for assessment of the petroleum potential offshore north-east greenland (greenland sea), and based on available outcrop samples the potential seemed surprisingly poor. three cores with nearly 100% recovery and a total length of more than 600 m were fig. 2. aerial view of kilen, a key area for understanding the geology of the wandel sea basin with exposures of triassic to cretaceous sediments. kilen is approximately 10 × 30 km large and is surrounded by glaciers except to the south-east that faces the northeast water polynya. view towards the south. for location see fig. 1. 61 collected, covering the full stratigraphy from the oxfordian to the ryazanian (blokelv-1, rødryggen-1 and brorson halvø-1). more than 600 samples were analysed for petroleum potential and numerous biological marker and isotopic analyses carried out. integrated with detailed stratigraphic and sedimentological studies they provide a unique insight into the character and development of the succession. based on these studies, it has been demonstrated that the petroleum potential is 2–3 times higher than expected from initial data and that the succession, contrary to previous beliefs, is complete with only local hiatuses in specific tectonic settings. core drilling has also been carried out in unexposed intervals of the cretaceous (store koldewey-1, nanok-1), and in the triassic (dunken-1, 2) and permian successions (kim fjelde-1) of peary land. traditional onshore field work focused on solving predefined problems related to the overall key risks and uncertainties listed in the foregoing. each season, individual field teams dedicated most of their efforts to one of the various projects (table 1), while also contributing to other projects where needed. for instance, although both sand provenance and uplift studies were carried out with dedicated field teams, all teams collected samples for mineralogical analysis, apatite fission track analysis (afta) and vitrinite reflectance analysis (ro), ensuring proper geographical and stratigraphic coverage. a small degree of overlap between some projects exists. for instance, the store koldewey study was concentrated in a restricted geographical area, but included deposits ranging from caledonian basement to pleistocene mud and elements of uplift studies, sand provenance, etc. as integral parts of the study. in contrast, the cretaceous study concerned a specific, poorly known, portion of the stratigraphic column with scattered outcrops over a huge area (including store koldewey). the study was divided into sub-projects such as the erection of a formal and consistent lithostratigraphy, an integrated biostratigraphy that included macrofossils, microfossils (palynomorphs), reservoir model studies, petroleum source rock studies and more. published research generally, all results produced during the course of the project are subject to a five-year confidentiality clause that must expire before geus can publish detailed results. however, general observations of little or no relevance for the proprietary studies have been published and more are under way. these include studies of igneous rocks (larsen et al. 2013, in press), coal deposits (bojesen-koefoed et al. 2012; petersen et al. 2013), palaeogene sediments (nøhr-hansen et al. 2011), stratigraphy and palaeoenvironments (alsen & mutterlose 2009; nøhr-hansen 2012; pauly et al. 2012a, b, 2013) and integrated studies (fyhn et al. 2012). a large number of pathe triassic of the wandel sea basin pending also highly relevant for barents sea geology the jurassic–cretaceous of the pending also highly relevant for barents sea geology wandel sea basin the tectonic evolution of the pending also highly relevant for barents sea geology wandel sea basin northern east greenland uplift history pending extension of previous uplift studies to 83°n and northwards – focus on the wandel sea basin table 1. overview of completed and proposed projects project title initiated completed comments data package, blokelv-1 core, oct. 2008 dec. 2009 oxfordian–volgian portion of u. jurassic – l. cretaceous source jameson land rock succession. comprehensive sedimentological, stratigraphic and source rock study provenance study of possible oct. 2008 dec. 2010 based on zircon ages and provenance sensitive minerals. reservoir sandstone units in east comprehensive study of potential reservoir sandstone units, and north-east greenland >200 samples data package, rødryggen-1 core, mar. 2010 dec. 2010 kimmeridgian–ryazanian portion of u. jurassic – l. cretaceous source wollaston forland rock succession. comprehensive sedimentological, stratigraphical and source rock study north-east greenland uplift history mar. 2010 nov. 2011 extension of initial uplift study to 78°n phase ii the cretaceous of north-east jun. 2010 jun. 2012 comprehensive study of the entire exposed and mostly poorly greenland known cretaceous succession geology of store koldewey dec. 2010 mar. 2012 important area for understanding the danmarkshavn basin. gis compilation, including core-hole data data package, brorson halvø-1 core, apr. 2011 dec. 2011 kimmeridgian–ryazanian portion of u. jurassic – l. cretaceous wollaston forland source rock succession in different tectonic settings. comprehensive sedimentological, stratigraphic and source rock study data package, nanok-1 core, nov. 2011 nov. 2012 poorly exposed cenomanian–turonian succession hold with hope comprehensive sedimentological, stratigraphic and source rock study 6262 pers by geus scientists are expected in the coming years as confidentiality expires. impact on offshore petroleum exploration and perspectives for the future a better understanding of numerous critical parameters has been achieved, and many of the results have had a strong influence on the development of offshore exploration models. in particular, based on the onshore work a better understanding of source rock distribution, thickness and quality as well as the timing and magnitude of subsidence and uplift have been obtained. in addition, a detailed understanding of the cretaceous succession will be very important for evaluating future prospects and for correlation when offshore core holes and wildcat wells are eventually drilled. in recent years, the focus of onshore field activities has shifted northwards, where the deposits of the wandel sea basin (fig. 1) are important for the understanding of the geology of the northernmost portion of the licence area off north-east greenland. the northern region is also important for understanding the geology of the barents sea shelf, which is an area under intense exploration. hence, a gradual shift of focus towards the conjugate margin is expected to take place in the coming years. the nature of the collaboration with the industry in north-east greenland must be expected to change after allocation of licences. companies holding licences may change their focus whereas others not holding licences are expected to downgrade their interests in the region. however, geus sees an obvious interest in a continued collaboration, since it provides a platform for studies that would otherwise be beyond the economic capacity of the institution, but which are nonetheless important elements of the raison d’être of a geological research institution. references alsen, p. & mutterlose, j. 2009: the early cretaceous of north-east greenland: a crossroads of belemnite migration. palaeogeography, palaeoclimatology, palaeoecology 280, 168–182. bojesen-koefoed, j.a., bjerager, m. & piasecki, s. 2009: shallow core drilling and petroleum geology related field work in north-east greenland 2008. geological survey of denmark and greenland bulletin 17, 53–56. bojesen-koefoed, j.a., kalkreuth, w., petersen, h.i. & piasecki, s. 2012: a remote coal deposit revisited: middle jurassic coals at kulhøj, western germania land, northeast greenland. international journal of coal geology 98, 50–61. christiansen, f.g. 2011: greenland petroleum exploration: history, breakthroughs in understanding and future challenges. memoir of the geological society (london) 35, 647–661. christiansen, f.g., gautier, d.l., stemmerik, l., bidstrup, t., bojesenkoefoed, j.a. & sørensen, k. 2006: petroleum resource potential of the east greenland shelf. aapg hedberg research conference on understanding world oil resources, colorado springs, colorado, usa, 12–17 november 2006. extended abstract, poster. fyhn, m.b.w., rasmussen, t.m., dahl-jensen, t., weng, w.l., bojesenkoefoed, j.a. & nielsen, t. 2012: geological assessment of the east greenland margin. geological survey of denmark and greenland bulletin 26, 61–64. gautier, d.l. 2007: assessment of undiscovered oil and gas resources of the east greenland rift basins province. u.s. geological survey fact sheet 2007-3077, 4 pp. gautier, d.l., stemmerik, l., christiansen, f.g., sørensen, k., bidstrup, t., bojesen-koefoed, j.a., bird, k.j., charpentier, r.r., houseknecht, d.w., klett, t.r., schenck, c.j. & tennyson, m.e. 2011: assessment of ne greenland: prototype for development of circum-arctic resource appraisal methodology. memoir of the geological society (london) 35, 663–672. hamann, n.e., whittaker, r.c. & stemmerik, l. 2005: geological development of the northeast greenland shelf. in: doré, a.g. & vining, a.a. (eds): petroleum geology: north-west europe and global perspectives. proceedings of the 6th petroleum geology conference, 887–902. london: geological society. larsen, l.m., pedersen, a.k., sørensen, e.v., watt, w.s. & duncan, r.a. 2013: stratigraphy and age of the eocene igtertivâ formation basalts, alkaline pebbles and sediments of the kap dalton group in the graben at kap dalton, east greenland. bulletin of the geological society of denmark 61, 1–18. larsen, l.m., pedersen, a.k., tegner, c. & duncan, r.a. in press: eocene to miocene igneous activity in ne greenland: northward younging of magmatism along the east greenland margin. journal of the geological society (london), http://dx.doi.org/10.1144/jgs2013-118 nøhr-hansen, h. 2012: palynostratigraphy of the cretaceous – lower palaeogene sedimentary succession in the kangerlussuaq basin, southern east greenland. review of palaeobotany and palynology 178, 59–90. nøhr-hansen, h., nielsen, l.h., sheldon, e., hovikoski, j. and alsen, p. 2011: palaeogene deposits in north-east greenland. geological survey of denmark and greenland bulletin 23, 61–64. pauly, s., mutterlose, j. & alsen, p. 2012a: lower cretaceous (upper ryazanian–hauterivian) chronostratigraphy of high latitudes (northeast greenland). cretaceous research 34, 308–326. pauly, s., mutterlose, j. & alsen, p. 2012b: early cretaceous palaeoceanography of the greenland–norwegian seaway evidenced by calcareous nannofossils. marine micropaleontology 90–91, 72–85. pauly, s., mutterlose, j. & alsen, p. 2013: depositional environments of lower cretaceous (ryazanian–barremian) sediments from wollaston forland and kuhn ø, north-east greenland. bulletin of the geological society of denmark 61, 19–36. petersen, h.i, øverland, j.a., solbakk t., bojesen-koefoed, j.a. & bjerager, m. 2013: unusual resinite-rich coals found in northeastern greenland and along the norwegian coast: petrographic and geochemical composition. international journal of coal geology 109–110, 58–76. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: jbk@geus.dk. geological survey of denmark and greenland bulletin 33, 2015, 61-64 61 unique applied glaciology challenges of proglacial mining william colgan, henrik højmark thomsen and michele citterio th e glaciology group at the geological survey of denmark and greenland (geus) has a tradition of applied glaciology consulting for hydropower projects in greenland (weidick & th omsen 1982; braithwaite & olesen 1988; ahlstrøm et al. 2008). th is includes assessments for the hydropower plants now operating at ilulissat and nuuk (braithwaite & th omsen 1989; th omsen et al. 1989; 1993), as well as the outburst potential of ice-dammed lakes such as qorlortorsuup tasia (mayer & schuler 2005). several factors, including long term increases in global resource demand, increasing air temperatures and glacier retreat due to climate change, and improved mining and prospecting techniques may now improve the economic feasibility of mining in greenland (colgan & arenson 2013). given that over 80% of greenland is ice-covered, mining projects in greenland oft en occur in ‘proglacial’ settings, meaning adjacent to, or close to, an ice margin. th e isukasia, kvanefj eld, maarmorilik and malmbjerg prospects exemplify resource development in proglacial settings in greenland. th e glaciology group at geus has a growing interest in the applied glaciology aspects of proglacial mining projects, and was recently engaged to assess ice fl ow, meltwater runoff , and supraglacial road access for the malmbjerg prospect in east greenland (citterio et al. 2009), as well as provide expert commentary on the kumtor mine in kyrgyzstan (satke & galdini 2014). approximately 3.5 km2 of glacier overburden was removed at the kumtor mine between 1998 and 2014, creating the world’s largest open ice pit, in order to recover subglacial ore (fig. 1). th e open ice pit approved in 2013 for the isukasia (or ‘isua’) mine would be several times larger (fig. 2). applied glaciology addresses a number of unique geotechnical challenges associated with proglacial mining projects, here, we briefl y review four: supraglacial runoff , subglacial water fl ow, ice movement and supraglacial access roads. 2 km 78°16´e 41°55´n 41°50´n mine area (1998) ice margin (1977) mine area (2014) ice margin (1977) 78°9.5´e fig. 1. kumtor mine, kyrgyzstan. historic ice margins and contemporary mine areas overlaid on 1998 (left) and 2014 (right) landsat images. © 2015 geus. geological survey of denmark and greenland bulletin 33, 61–64. open access: www.geus.dk/publications/bull 6262 unique challenges it can be diffi cult to quantify and manage supraglacial meltwater runoff near the ice sheet margin. in southern greenland, the annual ice melt (or ‘ablation’) can exceed 8 m water equivalent per unit area per melt season (fausto et al. 2012). by comparison, the record annual rainfall on earth is c. 12 m water equivalent in meghalaya, india. in greenland, however, the annual meltwater runoff is concentrated during a relatively brief summer melt season. during the melt season, water-saturated snowpacks are susceptible to slushfl ows, which can damage infrastructure both on, and adjacent to, glaciers (smart et al. 2000). while surface mass-balance parameters generated by regional climate models can be calibrated with in situ data to assess the magnitude and spatial distribution of meltwater production (van as et al. 2014), estimates of the runoff that drains to any given proglacial site are very sensitive to the delineation of the supraglacial catchment areas (arenson & colgan 2015). uncertainty in the catchment delineation over the relatively fl at ice sheet can easily contribute to 50% uncertainty in runoff (rennermalm et al. 2013). recent work has also demonstrated that supraglacial streams regularly breach local topographic divides, making it diffi cult to justify the application of strictly elevation-dependent terrestrial water routing algorithms to the ice sheet (smith et al. 2015). manual delineation of supraglacial catchments from high-resolution imagery is a time-consuming, but more accurate, alternative (th omsen et al. 1989). it is also important to assess the subglacial water fl ow reaching any given proglacial site, as the vast majority of meltwater ‘runoff ’ produced on the surface of the ice sheet enters the enand subglacial networks prior to discharge at the ice sheet margin (smith et al. 2015). analogous to terrestrial groundwater fl ow, the enand subglacial hydrologic networks of the ice sheet have traditionally been conceptualised as saturated porous fl ow, whereby the enand subglacial water fl ow is governed by hydraulic potentiometric surfaces that can be predicted using ice geometry (lewis & smith 2009). recent observations, however, indicate that channelised subglacial drainage and non-trivial head fl uctuations extend tens of kilometres inland beneath the ice sheet (chandler et al. 2013). unlike groundwater fl ow through rock or sediment, the transmissivity of channelised fl ow in glacier ice can change rapidly, within hours or days, in response to both frictional melting along conduit walls and opening or closing due to viscous creep. as supraglacial lakes ice margin (c. 2010) proposed pit area movement (m/a) 10 100 greenland 3 km49°40´w49°50´ 65°10´n ice margin c. 2010 proposed pit area movement (m/a) 10 100 fig. 2. the potential ‘isua’ mine, greenland, approved in 2013. contemporary ice margins, proposed approximate pit area, and winter 2005/06 ice surface velocity vectors (joughin et al. 2010). the background is a 2014 landsat image. 200 100 0 operational year 0.0 distance inland (km) el ev at io n (m ) 0 0.5 1 1.5 2 2.5 200 100 0 60 50 40 30 20 10 0 operational year 10.0 a b distance inland (km) el ev at io n (m ) 0 0.5 1 1.5 2 2.5 ic e ve lo ci ty (m /y ) fig. 3. cross-sections of a glacier tongue with ice velocities (m/year). a: the glacier tongue has an undisturbed profile in operational year 0. b: after 10 years of excavation the ice wall has a gradient of 33%. the ice velocities have increased from 10–20 m/year to >60 m/year at the crown. the complete animation is available at www.williamcolgan.net/som/creng113 (colgan 2014). 63 can catastrophically release large volumes of meltwater (c. 109 l) into the subglacial network via crevasse hydrofracture (liang et al. 2012), there is a strong impetus to understand the preferential subglacial drainage routes in the vicinity of any given proglacial site. glacier ice movement can be so imperceptible over shorter time scales (hours to days) that glacier ice is commonly, but not strictly correctly, characterised as a solid. over longer time scales (years to decades), the true non-newtonian fl uid character and appreciable movement of ice becomes evident (colgan & arenson 2013). th e creep of glacier ice under gravitational stress can cause operational diffi culties due to ice movement beneath or against infrastructure, such as waste dumps and processing facilities (citterio et al. 2009). an exceptionally challenging task is to forecast ice fl ow into open ice pit excavations. due to the non-linear dependency of ice velocity on both ice thickness and surface gradient, the excavation of an open ice pit increases subsequent ice fl ow into the open ice pit (colgan 2014). perturbation of a natural glacier profi le into an artifi cial ice wall with a gradient of 33% can increase crown velocities by an factor of fi ve (fig. 3). open ice pits therefore require continuous excavation of substantial ice volumes to maintain pit geometry. over the lifecycle of a proglacial mine, this may require excavating several times the ice volume of an open ice pit itself. unlike conventional open pits in hard rock, the relative rapidity with which ice benches deform means that even temporary stoppages in ice excavation can adversely aff ect ice pit shape and fl ow (els 2012). th e establishment and maintenance of supraglacial access roads are oft en critical elements of proglacial mining projects (citterio et al. 2009). traversing the relatively subdued topography of a glacier can be preferable to traversing the relatively severe topography of proglacial areas with abundant rivers and erratics (fig. 4). in addition to potential crevasse hazards, supraglacial access roads can be compromised due to horizontal and vertical ice movement and diff erential surface ablation. while horizontal ice movement is readily observable by satellite (fig. 2), vertical ice movement can be heterogeneous over short distances, where compressive and extensional fl ow result in opposing vertical movements via thrustand slip-faulting along ice fractures (nye 1952). it is desirable to establish a metre-scale aggregate, supraglacial road bed followed by continual grading in order to minimise the infl uence of heterogeneous vertical ice movements. th is also improves road stability by keeping the ice beneath a road frozen throughout the year. however, active suppression of ablation beneath a supraglacial road results in perching of a road above the surrounding glacier surface (davis 1967). perching and consequent shoulder slumping hazards make roads that traverse glaciers even more sensitive to thermodynamic conditions than roads that traverse permafrost areas. additionally, as a consequence of glacier margin retreat, which can exceed 10 m per year, approach ramps at the transition from nonto ice-covered terrain require persistent maintenance on a weekly basis during the melt season (davis 1967). summary climate change will further exacerbate the unique applied glaciological challenges associated with proglacial mining described above. rising atmospheric temperatures are expected to increase the meltwater runoff from the ice sheet by a factor of fi ve by the end of the century (fettweis et al. 2013). th e probability of individual catastrophic supraglacial lake drainage events into the subglacial system is proportional to this summer melt intensity (liang et al. 2012). as ice rheology, i.e., the relation between stress and strain, is highly temperature dependent, the tremendous latent energy of meltwater can heat the ice and accelerate ice deformation (phillips et al. 2013). th e challenge of maintaining supraglacial road access increases with the length of the melt season fig. 4. the potential kerr-sulphurets-mitchell mine in canada, approved in 2014. proposed approximate mine area, contemporary ice margins and supraglacial access road. the background is a 2014 landsat image. the analogous features of the nearby brucejack prospect are also shown. brucejack kerr-sulphurets-mitchell ice margin (2012) 56°24´ 56°32´ 130°18´ 5 km130°4´e 56°16´ 56°24´n 6464 authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: wic@geus.dk and surface ablation (davis 1967). finally, with glaciers serving as highly visible indicators of climate change, proglacial mining projects oft en face exceptional public opposition in comparison to conventional hard rock mining projects (e.g. satke & galdini 2014). as described in this report, the geus glaciology group has both a history of providing commercial services for proglacial projects, and a familiarity with the unique challenges confronting future proglacial projects. th e glaciology group has carried out extensive research and monitoring programmes in greenland for more than sixty years. during this time, the glaciology group has developed a broad suite of glaciology-specifi c instrumentation, including autonomous climate stations that are capable of reporting numerous parameters via satellite link at hourly time intervals. th e ongoing programme for monitoring of the greenland ice sheet (promice), generates a large amount of climatology and glaciology data, as well as ensures a glaciology group presence for survey and maintenance operations around the ice sheet perimeter. when synergistic, the glaciology group partners with asiaq (greenland survey) on joint projects. in addition to serving as an advisory institute to the danish ministry of climate and energy, geus also serves as a contractually appointed advisory institute to the greenland ministry of mineral resources. references ahlstrøm, a., mottram, r., nielsen, c., reeh, n. & andersen, s. 2008: evaluation of the future hydropower potential at paakitsoq, ilulissat, w. greenland. danmarks og grønlands geologiske undersøgelse rapport 2008/37, 50 pp. arenson, l. and colgan, w. 2015: water management challenges associated with mining projects in greenland, 533–543. proceedings of mine water solutions in extreme environments 2015. vancouver, canada. braithwaite, r.j. & olesen, o. 1988: eff ect of glaciers on annual run-off , john dahl land, south greenland. journal of glaciology 34, 200–207. braithwaite, r.j. & th omsen, h.h. 1989: simulation of run-off from the greenland ice sheet for planning hydro-electric power, ilulissat/jakobshavn, west greenland. annals of glaciology 13, 12–15. chandler, d.m. et al. 2013: evolution of the subglacial drainage system beneath the greenland ice sheet revealed by tracers. nature geoscience 6, 195–198. citterio, m., mottram, r., larsen s.h. & ahlstrøm, a. 2009: glaciological investigations at the malmbjerg mining prospect, central east greenland. geological survey of denmark and greenland bulletin 17, 73–76. colgan, w. 2014: considering the ice excavation required to establish and maintain an open ice pit. journal of cold regions engineering 28, 04014003. colgan, w. & arenson, l.u. 2013: open-pit glacier ice excavation: brief review. journal of cold regions engineering 27, 223–243. davis, r.m. 1967: ice surface movement on the tuto ramp in north greenland. u.s. army cold regions research & engineering laboratory technical report 164, 24 pp. els, f. 2012: you want ice with that? centerra dives 18% aft er saying waste and ice will cut kumtor production by 200,000 oz. mining.com, 27 march 2012. http://www.mining.com/you-want-ice-with-that-centerra-dives-18-aft er-saying-waste-and-ice-will-cut-kumtor-production-by200000-oz. fausto, r.s., van as, d. & the promice project team 2012: ablation observations for 2008–2011 from the programme for monitoring of the greenland ice sheet (promice). geological survey of denmark and greenland bulletin 26, 73–76. fettweis, x., franco, b., tedesco, m., van angelen, j.h., lenaerts, j.t.m., van den broeke, m.r. & gallée, h. 2013: estimating the greenland ice sheet surface mass balance contribution to future sea level rise using the regional atmospheric climate model mar. th e cryosphere 7, 469–489. joughin, i. smith, b.e., howat, i.m. scambos, t. & moon, t. 2010: greenland fl ow variability from ice-sheet-wide velocity mapping. journal of glaciology 56, 415–430. lewis, s.m. & smith, l.c. 2009: hydrologic drainage of the greenland ice sheet. hydrological processes 23, 2004–2011. liang, y.l., colgan, w., lv, q. steff en, k., abdalati, w., stroeve, j., gallaher, d. & bayou, n. 2012: a decadal investigation of supraglacial lakes in west greenland using a fully automatic detection and tracking algorithm. remote sensing of environment 123, 127–138. mayer, c. & schuler, t.v. 2005: breaching of an ice dam at qorlortossuup tasia, south greenland. annals of glaciology 42, 297–302. nye, j.f. 1952: th e mechanics of glacier fl ow. journal of glaciology 2, 82–93. phillips, t., rajaram, h., colgan, w., steff en, k. & abdalati, w. 2013: evaluation of cryo-hydrologic warming as an explanation for increased ice velocities in the wet snow zone, sermeq avannarleq, west greenland. journal of geophysical research, earth surface 118, 1241–1256. rennermalm, a.k., smith, l.c., chu, v.w., box, j.e., forster, r.r., van den broeke, m.r., van as, d. & moustafa, s.e. 2013: evidence of meltwater retention within the greenland ice sheet. th e cryosphere 7, 1433–1445. satke, r. & galdini, f. 2014: ebrd’s environmental policy under scrutiny in kyrgyzstan. th e diplomat, 17 november 2014. http://thediplomat.com/2014/11/ebrds-environmental-policy-under-scrutiny-inkyrgyzstan smart, c.c., owens, i.f., lawson, w. & morris, a.l. 2000: exceptional ablation arising from rainfall-induced slushfl ows: brewster glacier, new zealand. hydrological processes 14, 1045–1052. smith, l.c. et al. 2015: effi cient meltwater drainage through supraglacial streams and rivers on the southwest greenland ice sheet. pnas 112, 1001–1006. th omsen, h.h., th orning, l. & olesen, o.b. 1989: applied glacier research for planning hydro-electric power, ilulissat/jakobshavn, west greenland. annals of glaciology 13, 257–261. th omsen, h.h., braithwaite, r.j., weidick, a. & olesen, o.b. 1993: evaluation of hydropower potential for possible future industrial use, nuuk area, west greenland. rapport grønlands geologiske undersøgelse 159, 59–62. van as, d. et al. 2014: increasing meltwater discharge from the nuuk region of the greenland ice sheet and implications for mass balance (1960–2012). journal of glaciology 60, 314–322. weidick, a. & th omsen, h.h. 1986: a decade of glacier investigations for utilisation of greenland hydropower. rapport grønlands geologiske undersøgelse 128, 157–169. http://www.mining.com/you-want-ice-with-that-centerra-dives-18-after-saying-waste-and-ice-will-cut-kumtor-production-by-200000-oz./ http://thediplomat.com/2014/11/ebrds-environmental-policy-under-scrutiny-in-kyrgyzstan/ << /ascii85encodepages false /allowtransparency false /autopositionepsfiles true /autorotatepages /pagebypage /binding /left /calgrayprofile (dot gain 15%) /calrgbprofile (srgb iec61966-2.1) /calcmykprofile (coated fogra27 \050iso 12647-2:2004\051) /srgbprofile (srgb iec61966-2.1) /cannotembedfontpolicy /warning /compatibilitylevel 1.7 /compressobjects /off /compresspages true /convertimagestoindexed true /passthroughjpegimages 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/romandefault /preserveediting true /untaggedcmykhandling /leaveuntagged /untaggedrgbhandling /leaveuntagged /usedocumentbleed false >> ] /syntheticboldness 1.000000 >> setdistillerparams << /hwresolution [2400 2400] /pagesize [595.276 841.890] >> setpagedevice geological survey of denmark and greenland bulletin 33, 2015, 65-68 65© 2015 geus. geological survey of denmark and greenland bulletin 33, 65–68. open access: www.geus.dk/publications/bull observed melt-season snowpack evolution on the greenland ice sheet charalampos charalampidis and dirk van as due to recent warm and record-warm summers in greenland (nghiem et al. 2012), the melt of the ice-sheet surface and the subsequent runoff are increasing (shepherd et al. 2012). about 84% of the mass loss from the greenland ice sheet between 2009 and 2012 resulted from increased surface runoff (enderlin et al. 2014). th e largest melt occurs in the ablation zone, the low marginal area of the ice sheet (van as et al. 2014), where melt exceeds wintertime accumulation and bare ice is thus exposed during each melt season. in the higher regions of the ice sheet (i.e. the accumulation area), melt is limited and the snow cover persists throughout the year. it is in the vast latter area that models struggle to calculate certain mass fl uxes with accuracy. a better understanding of processes such as meltwater percolation and refreezing in snow and fi rn is crucial for more accurate greenland icesheet mass-budget estimates (van angelen et al. 2013). in may 2012, the fi eld campaign ‘snow processes in the lower accumulation zone’ was organised by the geological survey of denmark and greenland (geus) at the kan_u automatic weather station (67°0´0˝n, 47°1́ 1˝w; 1840 m above sea level), which delivers data to the programme for monitoring of the greenland ice sheet (promice; van as et al. 2013) and is one of the few weather stations located in the lower accumulation area of greenland (fig. 1, inset). during the expedition, we installed thermistor strings, fi rn compaction monitors and a snowpack analyser; we drilled fi rn cores, performed fi rn radar measurements, gathered meteorological data, dug snow pits and performed dye-tracing experiments. one important objective of the campaign was to understand the thermal variability in the snowpack during the melt season by monitoring with high-precision temperature probes (campbell scientifi c temperature probe, model 107; accuracy: better than ± 0.4°c over the range –24 to 48°c). six temperature probes were installed in the snowpack of the previous winter at depths of 0.05, 0.10, 0.20, 0.30, 0.40 and 0.70 m below the surface (fig. 1). th e data from the probes were stored at 30-minute intervals on data loggers, which also triggered additional measurements of radiation-shielded air temperature at 1.10 m, surface albedo and surface-height change due to accumulation and ablation. emitted longwave radiation was also recorded to be able to calculate the surface temperature assuming snow to be a black-body radiator. th e vertical position of the probes relative to the surface, which changes due to ablation and accumulation, was determined by the sonic ranger measurements. recorded temperatures aft er the probes surfaced were discarded. th e relatively shallow snowpack (0.70–0.80 m) was on top of fi rn of density ρ >500 kg m-3 which had accumulated in the previous years (fig. 1). in the upper fi rn we found ice lenses (ρ >800 kg m–3) several metres thick. within the snowpack, two thin ice layers were present, one at 0.30 m greenland fig. 1. after the installation of the temperature probes in may 2012 (thin black cables) at the location of kan_u on the greenland ice sheet. the thick grey cables are thermistor strings drilled into the firn. the inset map shows the location of the study area. 6666 and one at c. 0.50 m below the surface, both about 0.01 m thick. th e average density of the snow was determined to be roughly 360 kg m–3, yielding an accumulation of 0.25 ± 0.08 m water equivalent (w.e.) since the summer 2011 (charalampidis et al. 2015). below, we present observations from the period 02 may to 23 july and interpret the atmosphere–surface interaction and its impact on the subsurface snow layers, with the goal to quantify refreezing in the greenland accumulation area. atmosphere–snow interaction th e observations reveal a strong similarity between the nearsurface air temperature and the snow-surface temperature with changes of air temparature lagging on average 30–40 minutes behind. typically, the air remained warmer than the snow surface (fig. 2), implying a prevailing stable stratifi cation of the near-surface air. at night, the diff erence was larger (1.6–1.9°c) due to the reduced sunlight and subsequent cooling of the surface forced by longwave radiation. during the day, the temperature diff erence was smaller (0.6–1.2°c) primarily due to solar radiation heating the surface and reducing atmospheric stability. understandably, when air temperature exceeds 0°c, the temperature diff erence can be larger since the surface cannot exceed the melting point. on 6 may, overcast conditions and atmospheric stillness caused the increase of air temperature above +6°c (fig. 3a). th ese were the highest temperatures during the observational period and similar temperatures occurred also on two days in june (3 and 18). with the exception of 6 may, the air temperature remained negative until the last week of may. during this period, the temperature of the upper 0.20 m of snow followed a pronounced diurnal cycle, which at 0.20 m lagged about 10 hours behind the variations in surface temperature (fig. 3b), signifying the low thermal conductivity of snow. at 1840 m above sea level, the ice sheet generally experiences low melt rates. when melt occurs, it displays a diurnal cycle following air temperature. a diurnal cycle of positive air temperatures occurred fi rst on 27 may, marking the beginning of the melt season (fig. 3a). th e surface ablated in response to the warm conditions, while the snow temperatures revealed the distinct progression of a warming and thus wetting front moving vertically through the snowpack (fig. 3b). th e temperature at 0.70 m depth was aff ected by this 42 hours aft er surface melt initiated, i.e. an average warming front progression of only c. 17 mm h–1. th e entire snowpack became temperate aft er six days of ablation. th e slow progression of the warming front indicates a concurrent heterogeneous meltwater infi ltration to the fi rn below (humphrey et al. 2012). during the period 8–12 june, sub-freezing air temperatures occurred again (fig. 3a) and melting ceased. while the upper part of the snowpack remained close to 0°c, possibly containing liquid water, the deeper levels (0.4 m and below) cooled as heat was conducted downwards into the colder fi rn. melt resumed on 13 june, aff ecting snow temperatures at 0.7 m aft er 30 hours (fig. 3b), which is faster than in the previous melt period due to the reduced measurement depth, and changed snow properties. th ereaft er, the snow remained at the melting point until it ablated completely on 11 july. between 13 june and 11 july, there were fi ve occasions when the diurnal air/surface temperature cycle was interrupted by periods with warm night-time conditions resulting in enhanced ablation (fig. 3a). most notably, during the warm week of 8–14 july when the whole greenland icesheet surface area was reported to melt (nghiem et al. 2012), the air temperature at kan_u remained above +2°c for six days and melt was large. at the same time, the ‘watson river’, which drains this section of the ice sheet, experienced the highest discharge in 56 years, judging from the partial destruction of a 1956 bridge near the town of kangerlussuaq. snowpack evolution in may, the area received 0.12 m of fresh snow on top of the existing snowpack (fig. 4b) and the albedo remained at fresh snow values of 0.8–0.9 (fig. 4a). on 27 may, the surface began ablating and by the end of the day 0.05 m of the fresh snow had melted away. in the period until 8 june the average ablation rate was 0.02 m day–1, reducing albedo to c. 0.75, primarily due to snow metamorphosis. in principle, the energy needed to make temperate a uniform snowpack 0.7 m thick at –10°c is equivalent to the energy necessary for melting 0.04 m of snow at 0°c and a density of 360 kg m–3. th erefore, the generation of 15 mm of meltwater and its refreezing within the snowpack raises its temperature to 0°c. by the beginning of june when the entire snowpack had reached 0°c, the fi rst 0.15 m of snow (i.e. c. 50 mm of meltwater) had ablated. th is implies that approximately 70% of the meltwater either percolated deeper 1 3 5 7 9 11 13 15 17 19 21 23 time of the day (h utc) period: 2 may to 23 july 2012 (t a − t s ) av er ag e (° c ) −1 0 1 2 3 fig. 2. average temperature difference between air and surface during the day. the error bars show two standard deviations. solar zenith time is at 15:08 utc. 67 into the fi rn or was retained in liquid form in the snow by capillary forces. from the beginning of june onward, by lack of cold content, the snowpack was able to respond immediately to surface forcings (fig. 4b), and all percolating meltwater was routed toward the underlying fi rn. th e cold conditions and melt pause from 8 to 12 june were accompanied by snowfall resulting in 0.05 m of fresh snow accumulation (fig. 4b), thereby increasing the albedo above 0.8 (fig. 4a). melt resumed on 15 june with an average ablation rate of 0.03 m day–1 and with albedo dropping as low as 0.7, indicative of wet snow. small snowfall events also occurred in the beginning of july. during the warm days of 9 and 10 july the ablation rate exceeded 0.05 m day–1, removing the last of the 2011–2012 winter snowpack and revealing the underlying, water-saturated fi rn. consequently, the albedo dropped below 0.7, enhancing melt through the meltalbedo feedback (box et al. 2012). simulated refreezing rates a combination of temperature measurements and thermal conductivity simulations reveals the amount of refrozen water in the snow. a heat-conduction model was used to simulate the evolution of subsurface temperatures, using measured surface temperature and surface height change as input. th e model was run at temporal and spatial resolutions of 10 minutes and 0.10 m, respectively, and was re-initialised each day by measured temperature profi les at 00:00 utc. th e effective conductivity of the snow is a function of snow density (sturm et al. 1997) and the specifi c heat of snow depends on temperature (yen 1981). density profi les were initialised based on snow pit density measurements at the installation of the probes and were updated throughout the run taking refreezing into account. th e diff erence between the simulated and measured temperatures at the end of the day is a measure of the added latent heat during the day, and thus of the daily refreezing rates. th e simulation reveals that during the fi rst week of melt starting 27 may, refreezing occurred at all measurement depths within the snowpack (fig. 5a). th e peak refreezing occurred on 30 may and at depths below 0.50 m. during this period the total refreezing rate in the snowpack was comparable to the average melt rate of that fi rst period of melt (6 kg m–2 day–1; fig. 5b). th e subsequent cold content reduction and thinning snowpack resulted in low refreezing values below 0.4 kg m–2 day–1 at all depths. note that refreezing rates during the sub-freezing early period of our simulation are non-zero and large near the surface. it is possible that shortwave penetration in the snow plays a role or that our conduction model is fl awed in conditions of large temperature gradients in well-ventilated, low-density snow, which is valid for the start of the simulation period. however, in terms of total refreezing the early results add up to small values (fig. 5b). 0 4 2012 a 06 may 20 may 03 jun 17 jun 01 jul 15 jul −30 −25 −20 −15 −10 −5 0 te m p er at u re ( °c ) b air (1.10 m) surface 0.05 m 0.10 m 0.20 m 0.30 m 0.40 m 0.70 m te m p. ( °c ) fig. 3. observed temperatures of the near-surface atmosphere (a) and (sub)surface (b). 0.7 0.9 a lb ed o 2012 a in it ia l d ep th f ro m s u rf ac e (m ) b 06 may 20 may 03 jun 17 jun 01 jul 0.7 0.4 0.3 0.2 0.1 0.0 −28 −24 −20 −16 −12 −8 −4 0 (°c) fig. 4. observed surface albedo (a) and thermal evolution (b) of the snowpack. the dark red contour signifies 0°c. fig. 5. calculated refreezing rates in the snow at 0.1 m spacial resolution (a) and combined (b). x 10−3 x 10−3 0.10 m 0.20 m 0.30 m 0.40 m 0.50 m 0.60 m 0.70 m 0 0.2 0.4 0.6 0.8 1 2012 a 06 may 20 may 03 jun 17 jun 01 jul 15 jul 0 2 4 6 r ef r. ra te b total r ef re ez in g ra te ( 1 0 3 k g m − 2 d ay − 1 ) 6868 th e non-zero values of roughly 0.1 kg m–2 day–1 at greater depth are considered the uncertainty for the entire simulation period. during the cold period in june, the refreezing rates increased again to 3 kg m–2 day–1 (fig. 5b), which is an indication that liquid water was available, primarily at depths 0.20–0.30 m, while the required cold content was being supplied by the surface. th is method of refreezing requires liquid water retention in the snow matrix while cold content becomes available, as opposed to the refreezing of meltwater percolating into layers at sub-freezing temperatures. th e heat between 0.60–0.70 m that was conducted to depths below the seasonal snow layer increased the available cold content, thus when melt occurred again, refreezing was prominent at those depths (15 june; fig. 5a). as the average melt rate aft er 13 june was c. 8 mm w.e. day–1, the refrozen water in the snowpack was less than 10% of this amount, implying liquid water retention or the routing of meltwater to the layers below. overall, the simulated density increase within the snowpack was between 70–80 kg m–3 for most levels. meltwater refreezing is a positive component in the mass budget (mass storage; harper et al. 2012), although in a warming climate with more frequent extreme melt conditions, the larger meltwater fl uxes in the snow and fi rn may result in rapid reduction of pore volume (van angelen et al. 2013). th e large melt of 2012 at the elevation of kan_u was a result of both high atmospheric temperatures (bennartz et al. 2013) and a relatively low albedo from the exposure of the water-saturated fi rn aft er the early removal of the relatively thin winter snowpack (charalampidis et al. 2015). th e high ice content of the fi rn as found during the measurement campaign is an indication of intense percolation during previous years. th ese snow processes are still quite poorly represented in modelling eff orts, also due to the dependency of horizontal meltwater runoff on the ice layers formed by refreezing. our results illustrate that especially the melt-albedo feedback in relation to pore-volume reduction makes the lower accumulation area of the greenland ice sheet highly responsive in a warming climate. acknowledgements th e data presented in this paper were gathered in close collaboration with the greenland analogue project. we are grateful to our snow processes in the lower accumulation zone project partners horst machguth, mike macferrin, andreas mikkelsen, rickard pettersson, katrin lindbäck, alun hubbard and sam doyle. th is is a publication in the framework of the programme for monitoring of the greenland ice sheet (promice) and contribution number 63 of the nordic centre of excellence svali, ‘stability and variations of arctic land ice’, funded by the nordic toplevel research initiative (tri). references bennartz, r., shupe, m.d., turner, d.d., walden, v.p., steff en, k., cox, c.j. kulie, m.s. miller, n.b. & pettersen, c. 2013: july 2012 greenland melt extent enhanced by low-level liquid clouds. nature 496, 83–86. box, j.e., fettweis, x., stroeve, j.c., tedesco, m., hall d.k. & steff en, k. 2012: greenland ice sheet albedo feedback: thermodynamics and atmospheric drivers. th e cryosphere 6, 821–839. charalampidis, c. et al. 2015: changing surface-atmosphere energy exchange and refreezing capacity of the lower accumulation area, west greenland. th e cryosphere discussions 9, 2867–2913. enderlin, e.m., howat, i.m., jeong, s., noh, m.-j., van angelen, j.h. & van den broeke, m.r. 2014: an improved mass budget for the greenland ice sheet. geophysical research letters 41, 866–872. harper, j., humphrey, n., pfeff er, w.t., brown, j. & fettweis, x. 2012: greenland ice-sheet contribution to sea-level rise buff ered by meltwater storage in firn. nature 491, 240–243. humphrey, n.f., harper, j.t. & pfeff er, w.t. 2012: th ermal tracking of meltwater retention in greenland’s accumulation area. journal of geophysical research: earth surface 117, f01010. nghiem, s.v., hall, d.k., mote, t.l., tedesco, m., albert, m.r., keegan, k., shuman, c.a., digirolamo, n.e. & neumann, g. 2012: th e extreme melt across the greenland ice sheet in 2012. geophysical research letters 39, l20502. shepherd, a. et al. 2012: a reconciled estimate of ice-sheet mass balance. science 338, 1183–1189. sturm, m., holmgren, j., könig, m. & morris, k. 1997: th e thermal conductivity of seasonal snow. journal of glaciology 43, 26–41. van angelen, j.h., lenaerts, j.t.m., van den broeke, m.r., fettweis, x. & meijgaard, e. 2013: rapid loss of fi rn pore space accelerates 21st century greenland mass loss. geophysical research letters 40, 2109–2113. van as, d., fausto, r.s., colgan, w.t., box, j.e. & the promice project team 2013: darkening of the greenland ice sheet due to the melt-albedo feedback observed at the promice weather stations. geological survey of denmark and greenland bulletin 28, 69–72. van as, d. et al. 2014: increasing meltwater discharge from the nuuk region of the greenland ice sheet and implications for mass balance (1960–2012). journal of glaciology 60, 314–322. yen, y.c. 1981: review of thermal properties of snow, ice and sea ice. crrel report 81–10, 27 pp. hanover, new hampshire: us army corps of engineerscold regions research and engineering laboratory. authors’ addresses c.c.* & d.v.a., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: cc@geus.dk *also at: department of earth sciences, uppsala university, villavägen 16, 752 36 uppsala, sweden. geological survey of denmark and greenland bulletin 41, 2018, 79-82 79 the greenland ice sheet has experienced an average mass loss of 142 ± 49 gt/yr from 1992 to 2011 (shepherd et al. 2012), making it a significant contributor to sea-level rise. part of the icesheet mass loss is the result of increased dynamic response of outlet glaciers (rignot et al. 2011). the ice discharge from outlet glaciers can be quantified by coincident measurements of ice velocity and ice thickness (thomas et al. 2000; van den broeke et al. 2016). as part of the programme for monitoring of the greenland ice sheet (promice; ahlstrøm et al. 2008), three airborne surveys were carried out in 2007, 2011 and 2015, with the aim of measuring the changes in greenland ice-sheet thicknesses. the purpose of the airborne surveys was to collect data to assess the dynamic mass loss of the greenland ice sheet (andersen et al. 2015). here, we present these datasets of observations from ice-penetrating radar and airborne laser scanning, which, in combination, make us able to determine the ice thickness precisely. surface-elevation changes between surveys are also presented, although we do not provide an in-depth scientific interpretation of these. instrumentation all three surveys were conducted using the same air greenland/norlandair de havilland dhc-6 twin otter aircraft, currently registered as tf-pof. this twin otter has been modified in such a way that part of the fuselage can be removed in the rear cargo hole providing an unobstructed view of the surface below the aircraft when airborne. the precise position of the aircraft (and instruments) is tracked by three geodetic dual-frequency gps receivers each connected to one of two gps antennas mounted on top of the aircraft. the orientation of the instruments is monitored by an inertial navigation system (ins). the primary ins is of the type honeywell h-764g. during the last two flights, we also installed a back-up ins of the type oxts inertial+2. for measuring snowor ice-surface elevations, a near infrared, airborne laser scanner (als; forsberg et al. 2001) was mounted in the rear cargo hole, alongside the inss. the als flown on the twin otter in 2007 was of the type riegl lms-q140i-60, which was upgraded to a riegl lms-q240i in 2011 and 2015. in 2007 and 2011, a 60 mhz coherent ice-penetrating radar, developed at the technical university of denmark (dtu), was also mounted to measure bedrock topography (christensen et al. 2000). survey design the survey flight path was designed as a polygon to encircle the entire greenland ice sheet where the surface of the ice circum-greenland, ice-thickness measurements collected during promice airborne surveys in 2007, 2011 and 2015 louise sandberg sørensen, sebastian b. simonsen, rené forsberg, lars stenseng, henriette skourup, steen savstrup kristensen and william colgan 2 km b 800 m ice surface bedrock transmit pulse 70.35 70.34 70.33 70.32 70.31 70.30 30.920 30.910 30.900 30.890 2400 2392 2384 2376 2368 2360 2352 2344 a la tit ud e (n or th ) longitude (west) el ev at io n (m ) fig. 1. a: example of full-resolution versus reduced-resolution (circles) airborne laser scanner (als) data. b: example of radargramme with a clear bedrock reflector. © 2018 geus. geological survey of denmark and greenland bulletin 41, 79–82. open access: www.geus.dk/bulletin http://www.geus.dk/bulletin 8080 is at an elevation of c. 1700 m above sea level, as well as to include survey lines over the centerline of several main outlet glaciers. the surveys have been carried out at four-year intervals (2007, 2011 and 2015). the planned flight path in 2007 left a data gap on the east coast (from c. 72n to c. 74n) which was bridged during the 2011 and 2015 surveys. all three surveys were planned to be carried out in august, as this timing represents the end of the melt season and ensures that the changes observed in surface elevations are not affected by individual accumulation events. due to bad weather conditions in august 2015, half of the survey (the part from constable pynt in east greenland clockwise to kangerlussuaq in west greenland) was carried out in october. the late acquisition of these data thus results in a potential bias of individual accumulation events due to snowfall in this dataset compared to the surveys in 2007 and 2011. as the flight path from 2011 was repeated in 2015, and since the bedrock elevation is not expected to change within this time frame, it was decided not to utilise the ice-penetrating radar on the last survey in 2015. surface-elevation data the als operates in the near-infrared wavelength band, which is reflected from the snow or ice surface. this means that data can only be acquired during periods without clouds or fog below the aircraft. the sampling frequency of the als instrument is 10 khz, resulting in 40 across-track scan lines per second. each of these scan lines consists of 250 individual elevation measurements on-ground. the scan angle of 60° and the typical flight height of c. 300 m result in a swath width on ground of c. 300 m with c. 1 m resolution. the processing of the als data combines the raw als data with the positioning data from the gps and altitude data from the ins. post-processing of the data includes visual inspection to filter out laser reflections from clouds. the positional uncertainty in both latitude and longitude is estimated to be ± 1 m, while the elevation uncertainty is estimated from track cross-over differences to be ± 0.05 to 0.1 m over flat surfaces. to reduce the file size and to create a dataset which is more comparable to the resolution of the bedrock data, the full-resolution als data have been reduced to a spatial resolution of c. 100 m. this has been done through simple averag30°w50°w 60°n 65°n 70°n 75°n 80°n 80°n 30°w50°w 60°n 65°n 70°n 75°n 80°n 80°n n –500 –250 0 250 500 750 1000 1250 1500 bedrock elevation (m) 3000 2500 2000 1500 1000 500 surface elevation (m) 2011 500 km 2011 500 km fig. 2. a: surface elevations along the promice circum-greenland flights in 2011. b: bedrock elevations along the promice circum-greenland flights in 2011. 81 ing of available height measurements along and across track. an example of full-resolution versus reduced-resolution data is shown in fig. 1a. the data are compiled in one file per year (als_yyyy.ave) and can be downloaded from http:// promice.dk/downloadairborne.html. as an example, the elevations from 2011 are shown in fig. 2a. bedrock-elevation data the ice-penetrating data acquisition consists of transmitting pulses at a pulse repetition frequency of 10 khz (i.e. sampling in the flight direction) and sampling the returned echo at 75 mhz, which results in 4096 samples per transmitted pulse. while internal scattering masks the desired echo, reflection and absorption within the ice sheet reduce the strength of the returned echo. substantial processing is therefore carried out to produce a radargramme that enhances the detection of the echo from the bottom of the ice-sheet. a semi-automatic layer detection program is used to digitalise the surface and bedrock layers individually. in some areas, primarily near the ice margin in south greenland, the radar was not able to detect the bedrock due to heavily crevassed ice or water present within the ice. figure 1b shows a good example of a radargramme where a bottom echo was obtained. based on radar system setup, vertical uncertainty in radar-derived icesheet bed elevation is estimated to be ± 35 m, which is confirmed by the cross-over differences between the two surveys. the data are compiled in one file per year (ars_yyyy. ave), which is also available for download from http://promice.dk/downloadairborne.html. as an example, the bedrock elevations from 2011 are shown in fig. 2b. surface-elevation changes having three surveys of surface elevations spanning eight years enables us to derive and analyse surface elevation changes along the flight lines. in fig. 3, we show the mean annual surface-elevation changes between august 2007 and august/october 2015. the map was generated by computing height differences between any points in the two (reduced resolution) datasets for the two years. height differences are computed only if the points are located not more than 200 m apart. by knowing the exact date of the survey, the rate of surface-elevation change can be computed. in the map in fig. 3, the part that was only flown in 2007 is plotted with black, while the parts only surveyed in 2015 are shown in grey. there are some clearly visible gaps: one leg of the flight line is missing in north-eastern greenland from nioghalvfjerdsfjorden to hagen bræ and similarly and a part of the line is also absent south of jakobshavn isbræ. the gap in the north is caused by gaps in the 2015 dataset due to time and weather constraints. the gap south of jakobshavn isbræ is due to cloud cover in 2007. figure 3 shows that the mean annual elevation changes in the period 2007–2015 is clearly dominated by thinning with some main outlet glaciers such as jakobshavn isbræ and kangerlussuaq gletscher thinning rapidly. only a few places along the flight line are associated with thickening, e.g. at storstrømmen. the sections of the flight path in the southeastern parts that actually show modest thickening might be a result of accumulation since these parts of the 2015 survey were mapped in october after some snowfall in the area. elevation change data, such as presented here, are scientifically very valuable e.g. to validate satellite data and ice-sheet models. furthermore, the data presented here represens an important supplement to the heights and height differences 2 1 2.0 1.5 1.0 0.5 0.0 –0.5 –1.0 –1.5 –2.0 surface elevation change (m/yr) 30°w50°w 60°n 65°n 70°n 75°n 80°n 2007–2015 500 km 80°n 1 4 7 6 3 5 fig. 3. mean annual surface-elevation changes between 2007 and 2015 along the promice circum-greenland flight-paths. the part of the flight track for which only 2007 data are available is indicated in black, while 2015-only is indicated in grey. 1: nioghalvfjerdsfjorden. 2: storstrømmen. 3: constable pynt. 4: kangerlussuaq gletscher. 5: kangerlussuaq. 6: jakobshavn isbræ. 7: hagen bræ. http://promice.dk/downloadairborne.html http://promice.dk/downloadairborne.html http://promice.dk/downloadairborne.html http://promice.dk/downloadairborne.html 8282 available from the nasa operation icebridge field surveys (krabill et al. 2009; krabill 2014) as the flight lines cover different areas, and also our measurements are made at the end of the melt season while operation icebridge data are collected mainly in the spring. comparison to bedmachine v3 bedrock elevations the bedrock elevation dataset described above also represents a valuable legacy dataset that can be used by a wider scientific community. knowledge of bedrock elevations in greenland is essential in, e.g. ice-discharge studies and ice-sheet modelling. one widely used bedrock topography model is the one available in bedmachine v3 (morlighem et al. 2017) which is based on the conservation of mass and constrained by available measurements. the bedmachine v3 model is provided together with an error map, which shows how the error increases with increasing distance to measurement points. to evaluate whether the promice dataset can potentially contribute to an improvement of the bedmachine model in the future, we have extracted the bedmachine error values for all the 2007 and 2011 bedrock elevations in the promice datasets. the two corresponding histograms in fig. 4 show that in c. 50% of the data locations the error in the bedmachine v3 model is greater than 100 m, indicating that the promice dataset with an uncertainty of ± 35 m could indeed contribute positively to a future, improved version of the model. it may also be noted that only 25% of the bedmachine data are related with similar or lower errors than the promice dataset. acknowledgements this is a publication in the framework of the programme for monitoring of the greenland ice sheet (promice) – a danish government initiative funded through the danish cooperation for environment in the arctic (dancea). references ahlstrøm, a. & the promice team 2008: a new programme for monitoring the mass loss of the greenland ice sheet. geological survey of denmark and greenland bulletin 15, 61–64. andersen, m. et al. 2015: basin-scale partitioning of greenland ice sheet mass balance components (2007–2011). earth and planetary science letters 409, 89–95. christensen, e.l., reeh, n., forsberg, r., jørgensen, j.h., skou, n. & woelders, k. 2000: a low-cost glacier-mapping system. journal of glaciology 46, 531–537. forsberg, r., keller, k. & jacobsen, s.m. 2001: laser monitoring of ice elevations and sea-ice thickness in greenland. international archives of photogrammetry and remote sensing 34, 163–168. krabill, w.b. 2014: icebridge atm l2 icessn elevation, slope, and roughness. boulder, colorado, usa. http://nsidc.org/data/ilatm2. html (nasa distributed active archive center at the national snow and ice data center). krabill, w.b. et al. 2009: operation ice bridge =verview and results from aircraft laser altimetry. american geophysical union, fall meeting 14–18 december 2009. san francisco: abstract 3 pp. morlighem, m. et al. 2017: bedmachine v3: complete bed topography and ocean bathymetry mapping of greenland. from multibeam echo sounding combined with mass conservation. geophysical research letters 44, 11051–11061, http://dx.doi.org/10.1002/2017gl074954 rignot, e., velicogna, i., van den broeke, m.r., monaghan, a. & lenaerts, j.t. 2011: acceleration of the contribution of the greenland and antarctic ice sheets to sea level rise. geophysical research letters 38, l05503. shepherd, a. et al. 2012: a reconciled estimate of ice-sheet mass balance. science 338, 1183–1189. thomas, r.r., akins, t., csatho, b., fahnestock, m., gogineni, p., kim, c. & sonntag, j. 2000: mass balance of the greenland ice sheet at high elevations. science 289, 426–428. van den broeke, m.r., enderlin, e.m., howat, i.m., kuipers munneke, p., noël, b.p.y., van de berg, w.j., van meijgaard, e. & wouters, b. 2016: on the recent contribution of the greenland ice sheet to sea level change. the cryosphere 10, 1933–1946. http://dx.doi.org/10.5194/tc10-1933-2016 c ou nt s 0 100 200 300 400 500 2011 2007 error (m) 25 000 20 000 15 000 10 000 5000 0 30–40 m 40–100 m >100 m <30 m 24.4 8.8 50.0 16.8 fig. 4. histograms showing the errors of the bedmachine bed topography grid in all the points where promice bedrock elevation data are available. the grey area shows the <35 m interval (uncertainty in the bedrock data). the pie chart shows to what extent the bedmachine model error is 0–30 m, 30–40 m, 40–100 m and more than 100 m. authors’s addresses l.s.s., s.b.s., r.f., l.s. & h.s., technical university of denmark, dtu space, geodynamics department, dk-2800 kongens lyngby, denmark. e-mail: slss@space.dtu.dk. s.s.k., technical university of denmark, dtu space, microwave & remote sensing department, dk-2800 kongens lyngby, denmark. w.c., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. http://nsidc.org/data/ilatm2.html http://nsidc.org/data/ilatm2.html http://dx.doi.org/10.1002/2017gl074954 http://dx.doi.org/10.5194/tc-10-1933-2016 http://dx.doi.org/10.5194/tc-10-1933-2016 mailto:slss@space.dtu.dk geological survey of denmark and greenland bulletin 42, 2018, 1-5 petroleum geology of the upper jurassic – lower cretaceous of east and north-east greenland: blokelv-1 borehole, jameson land basin edited by jon r. ineson and jørgen a. bojesen-koefoed geological survey of denmark and greenland ministry of energy, utilities and climate geological survey of denmark and greenland bulletin 42 • 2018 22 geological survey of denmark and greenland bulletin 42 keywords jameson land, central east greenland, blokelv-1, borehole, petroleum geology cover recovery of the excellent blokelv-1 core, on which this bulletin is based, was thanks to the professional expertise of the drilling team (faxe kalk a/s, geus), depicted here erecting the drilling rig. photograph: annette ryge. frontispiece: facing page the flat-topped hills in the foreground are formed by the hareelev formation, the subject of this bulletin, viewed here in central jameson land with a backdrop to the east of the snow-clad mountains of liverpool land. photograph: jørgen bojesen-koefoed. chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, department of geoscience, aarhus university; minik rosing, geological museum, university of copenhagen; finn surlyk, department of geosciences and natural resource management, university of copenhagen scientific editors of this volume: jon ineson and jørgen bojesen-koefoed editorial secretary: jane holst referees (numbers refer to first page of reviewed article): morten smelror, no & william wimbledon, uk (15); michael larsen & finn surlyk, dk (39); jens therkelsen, dk & knut bjørlykke, no (65); erdim idiz & iain c. scotchmann, uk (85); jiri slama, cz & martin sønderholm, dk (115); christian tegner, dk & godfrey fitton, uk (127); andrew carter & andrew whitham, uk (133); finn surlyk, dk & rikke bruhn, no (149) illustrations: jette halskov and stefan sølberg layout and graphic productions: jacob lind bendtsen printers: rosendahls, søborg, denmark submission/acceptance dates of manuscripts: see end of individual articles printed: 28 december 2018 issn (print) 1604-8156 issn (online) 1904-4666 isbn (print) 978-87-7871-508-1 isbn (online) 978-87-7871-509-8 citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 42, 168 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark to buy bulletin in printed form please contact bogsalg@geus.dk and at www.geus.dk/bulletin42 (open access) © de nationale geologiske undersøgelser for danmark og grønland (geus), 2018 for the full text of the geus copyright clause, please refer to www.geus.dk/bulletin mailto:bogsalg@geus.dk http://www.geus.dk/bulletin42 http://www.geus.dk/bulletin 44 contents preface jørgen a. bojesen-koefoed ................................................................................................................................................................ 5 the upper jurassic blokelv-1 cored borehole in jameson land, east greenland – an introduction morten bjerager, stefan piasecki and jørgen a. bojesen-koefoed ............................................................................................... 7 biostratigraphy of the hareelv formation (upper jurassic) in the blokelv-1 core, jameson land, central east greenland peter alsen and stefan piasecki ....................................................................................................................................................... 15 sedimentology, geochemistry and reservoir properties of upper jurassic deep marine sediments (hareelv formation) in the blokelv-1 borehole, jameson land basin, east greenland morten bjerager, claus kjøller, mette olivarius, dan olsen and niels h. schovsbo .............................................................. 39 diagenesis of upper jurassic sandstones of the blokelv-1 core in the jameson land basin, east greenland mette olivarius, rikke weibel, niels h. schovsbo, dan olsen and claus kjøller ................................................................... 65 petroleum potential of the upper jurassic hareelv formation, jameson land, east greenland jørgen a. bojesen-koefoed, morten bjerager, h. peter nytoft, henrik i. petersen, stefan piasecki and anders pilgaard ............................................................................................................................................... 85 provenance of basinal sandstones in the upper jurassic hareelv formation, jameson land basin, east greenland mette olivarius, morten bjerager, nynke keulen, christian knudsen and thomas f. kokfelt ........................................... 115 igneous intrusions in the cored upper jurassic succession of the blokelv-1 borehole, jameson land basin, east greenland lotte melchior larsen ..................................................................................................................................................................... 127 burial and exhumation history of the jameson land basin, east greenland, estimated from thermochronological data from the blokelv-1 core paul f. green and peter japsen ...................................................................................................................................................... 133 late jurassic evolution of the jameson land basin, east greenland – implications of the blokelv-1 borehole morten bjerager, peter alsen, jørgen a. bojesen-koefoed, tove nielsen, stefan piasecki and anders pilgaard ............................................................................................................................................. 149 5 preface this bulletin presents a series of nine papers dealing with the succession of upper jurassic – lower cretaceous sedimentary rocks penetrated by the fully cored blokelv-1 borehole, drilled in western jameson land, central east greenland in august 2008. the borehole was drilled as the first of three boreholes that in combination were designed to provide full coverage of the upper jurassic – lower cretaceous petroleum source-rock succession in eastern greenland. the remaining two boreholes, rødryggen-1 and brorson halvø-1, were drilled on wollaston forland in 2009 and 2010, respectively, and the results from these boreholes will be published in a companion volume. the objectives of the drilling campaign were fulfilled, demonstrating that continuous sedimentation of oil-prone petroleum source rocks took place in eastern greenland over a period of c. 13 million years from the oxfordian to the ryazanian, with the blokelv-1 succession representing the older, oxfordian– volgian part of this interval. the drilling campaign was carried out as one of a number of projects within the framework of a multi-client collaborative programme between geus and a long list of petroleum companies entitled petroleum geological studies, services and data in east and northeast greenland. this collaboration was initiated in 2007 and is ongoing at the time of writing with more than 20 participant companies, a subset of which sponsored the studies presented herein; for contractual reasons, these companies cannot be named. the geus–industry collaboration was initiated in recognition of the need for new and better data on many aspects of the petroleum geology of eastern greenland prior to an anticipated licensing round of offshore north-east greenland. the circum-arctic resource appraisal (cara), undertaken by the united states geological survey (usgs), also played an important role in defining the priorities of the collaborative agreement by directing attention towards specific subjects in need of investigation. licensing rounds in 2012 and 2013 resulted in the award of five licences. based on the results of these activities in eastern greenland, a large number of scientific papers have been published since 2008, and more are expected as confidentiality clauses expire. this volume is, however, the first geus bulletin to be published as a direct consequence of the geus–industry collaboration. jørgen a. bojesen-koefoed © geus, 2018. geological survey of denmark and greenland bulletin 42, 5. available at: www.geus.dk/bulletin42 http://www.geus.dk/bulletin42 geological survey of denmark and greenland bulletin 1, 527-541 527 jurassic sedimentary rocks in sweden are restricted to skåne (scania), and adjacent offshore areas. they were deposited in areas separated by tectonic structural elements, each area reflecting different depositional and tectonic settings. the major areas are the ängelholm trough, the helsingborg area, the landskrona–kävlinge area, south-west skåne, central skåne, the vomb trough and the fyledalen fault zone (fig. 1). skåne is situated at the transition between the danish basin to the south-west and landmasses including the baltic shield to the north-east. therefore, small relative sea-level changes played a significant role in controlling the lateral facies distribution. the jurassic successions in skåne comprise sediments that accumulated in coastal plain to shallow shelf environments. the present distribution of jurassic deposits in skåne represents remnants of a once more extensive cover (norling et al. 1993). cretaceous–cenozoic inversion resulted in significant erosion of the deposits once deposited within the sorgenfrei–tornquist zone (norling & bergström 1987; erlström et al. 1997). lower jurassic deposits are by far the most widely preserved. they occur in troughs such as the ängelholm trough, the landskrona–kävlinge area, south-west skåne includthe jurassic of skåne, southern sweden anders ahlberg, ulf sivhed and mikael erlström in sweden, jurassic strata are restricted to skåne and adjacent offshore areas. jurassic sedimentary rocks predominantly comprise sandy to muddy siliciclastics, with subordinate coal beds and few carbonate-rich beds. during mesozoic times, block-faulting took place in the sorgenfrei– tornquist zone, a tectonic zone which transects skåne in a nw–se direction. the jurassic depositional environments in skåne were thus strongly influenced by uplift and downfaulting, and to some extent by volcanism. consequently, the sedimentary record reveals evidence of numerous transgressions, regressions and breaks in sedimentation. relative sea-level changes played a significant role in controlling the facies distribution, as deposition mainly took place in coastal plain to shallow shelf environments. the alluvial deposits in skåne include floodplain palaeosols, autochthonous coals, overbank sandstones, and stream channel pebbly sandstones. restricted marine strata comprise intertidal heteroliths with mixed freshwater and marine trace fossil assemblages, and intertidal delta distributary channel sandstones. shallow marine sediments encompass subtidal and shoreface sandstones with herringbone structures, and bioturbated mudstones with tempestite sandstones. offshore deposits typically comprise extensively bioturbated muddy sandstones. floral remains, palaeopedology, clay mineralogy and arenite maturity indicate a warm and humid climate in skåne throughout the jurassic, possibly with slightly increasing aridity towards the end of the period. most jurassic strata in skåne have been subjected to mild burial diagenesis, and the petroleum generative window has rarely been reached. keywords: skåne, southern sweden, fennoscandian border zone, danish basin, jurassic, lithostratigraphy, depositional environments, diagenesis a.a., department of geology, lund university, sölvegatan 13, s-223 62 lund, sweden. e-mail: anders.ahlberg@geol.lu.se u.s. & m.e., geological survey of sweden, kiliansgatan 10, s-223 50 lund, sweden. geological survey of denmark and greenland bulletin 1, 527–541 (2003) © geus, 2003 528 ing the höllviken graben, the hanö bay basin, the vomb trough and the fyledalen fault zone, and as erosional remnants on precambrian basement in central skåne. the rhaetian – lower jurassic succession is up to 250 m in thickness (sivhed 1984), whereas the middle and upper jurassic sediments are 75–300 m thick (norling 1972; norling et al. 1993). in skåne, most information on jurassic strata comes from the helsingborg area (norling 1972; sivhed & wikman 1986; norling & wikman 1990) and in the fyledalen fault zone, where relatively well-exposed sections occur (norling et al. 1993). in other areas, the development of the jurassic is incompletely known. results presented by vossmerbäumer (1969, 1970), rolle et al. (1979), ahlberg (1990, 1994), erlström et al. (1991, 1994), pieńkowski (1991a, b) and arndorff (1994) have added to the understanding of the depositional conditions and the relationship between the different areas of deposition. the aim of the present paper is to give an overview of the work hitherto done on the stratigraphy and sedimentology of the onshore jurassic strata in skåne, including previously unpublished discoveries and ideas. structural framework throughout the jurassic, block faulting played a significant role in skåne. the faulting was strongly influenced by major plate tectonic events, i.e., the break-up of pangea, the closing of tethys and the opening of the north atlantic (ziegler 1990). the early alpine tectonic phases in northwest europe resulted in the activation of the tornquist zone along lineaments established during the palaeozoic (norling & bergström 1987; erlström et al. 1997). the tornquist zone (fig. 1) is subdivided into the sorgenfrei–tornquist and teisseyere– tornquist zones (see discussions in liboriussen et al. 1987; norling & bergström 1987; eugeno-s working group 1988; berthelsen 1992; michelsen & nielsen 1993; thomas et al. 1993; thybo et al. 1994; erlström et al. 1997; vejbæk 1997). the sorgenfrei–tornquist zone links with the rønne graben and continues north-westwards through skåne to the northwestern part of jylland. the teisseyere–tornquist zone extends south-east from the rønne graben into poland. in general, tectonism actively controlled deposition and erosion in skåne throughout the jurassic, leaving bornholm skurup platform røn ne gra be n höllviken graben hanö baydenmark skåne romele block romeleåsen fault zone fyledalen fault zone kullen–ringsjön–andrarum fault zone colonus shale trough sweden 25 km sv ed al a fa ul t z on e ä h v c s f k l 14°e 56°n 55°n norway sweden denmark fig. 1. major tectonic elements of skåne, after erlström et al. (1997). stippled area índicates the sorgenfrei–tornquist zone, at the transition from the baltic shield in the north-east to the danish basin in the south-west. ä, ängelholm trough; c, central skåne; f, fyledalen fault zone; h, helsingborg area; k, kävlinge area; l, landskrona area; s, sw skåne; v, vomb trough. for additional outcrop and core locations, see sivhed (1984) and norling et al. (1993). different patterns of deposition and lacunas in individual tectonic blocks. major structural elements include the romele block, the höllviken graben, the colonus shale trough and the ängelholm trough. major faults include the kullen–ringsjön–andrarum fault zone, the fyledalen fault zone, the romeleåsen fault zone and the svedala fault zone (fig. 1; thomas et al. 1993; erlström et al. 1997). parts of the romele block, including the vomb trough (formed in the late cretaceous), were subjected to erosion and non-deposition during the late middle jurassic, while others persisted as depocentres. within the fyledalen fault zone, jurassic deposits are tilted or overturned due to uplift of the colonus shale trough (norling & bergström 1987; erlström et al. 1997). in the colonus shale trough, most triassic and jurassic deposits were removed by erosion associated with the late cretaceous – neogene inversion tectonics of the sorgenfrei–tornquist zone. a similar development has been observed in the rønne graben, which indicates the existence of a linked graben structure superimposed on the tornquist zone during much of the jurassic (vejbæk 1985; thomas & deeks 1994; vejbæk et al. 1994; erlström et al. 1997). depositional development in skåne, upper triassic (norian) fining-upwards cycles of coarse-grained, alluvial fan deposits grade upwards and basinwards into smectite-rich siltstones and mudstones, in which caliche nodules and halite pseudomorphs have been found (arndorff 1994). at the onset of the rhaetian, mineralogically mature, coal-bearing, kaolinite-rich siliciclastics were deposited, reflecting a regional change from warm semi-arid to warm and humid conditions, and an accompanying reduction in relief. the humidity was enhanced by the opening of interior seaways by rifting (manspeizer 1994). these climatic conditions dominated in skåne in the jurassic (hallam 1994). the lower jurassic deposits in skåne accumulated during a relatively quiet tectonic period, with a regional downwarp towards the depocentres of the danish basin (norling & bergström 1987). the succession in northwestern skåne is characterised by rhaetian–hettangian coastal plain to deltaic sediments succeeded by shallow marine deposits of sinemurian–aalenian age. in central skåne, rhaetian–hettangian sediments unconformably overlie deeply weathered precambrian gneiss. volcanic necks and associated pyroclastic deposits occur in central skåne (norin 1933); these have been dated to the early jurassic – middle jurassic transition (tralau 1968, 1973; bylund & halvorsen 1993). as a result of middle jurassic rifting in the north sea (ziegler 1990), tectonic unrest was transferred to skåne, which was subjected to uplift (norling & bergström 1987). marine influence ceased, and most of skåne was subjected to erosion which probably removed much of the lower jurassic succession. however, in certain areas (e.g. ängelholm trough and the fyledalen fault zone), deposition continued in coastal plain to shallow marine environments. during the late jurassic, new marine transgressions affected skåne (norling & bergström 1987). coastal and shallow marine sediments of oxfordian age were deposited, after a period of rapid facies changes in the late middle jurassic (norling 1981; norling & bergström 1987; erlström et al. 1997). after a period of relative quiescence, tectonic activity in skåne increased again at the jurassic–cretaceous transition, which led to reactivation (inversion) of pre-existing fault systems (erlström et al. 1997). in general, a warm and humid climate prevailed in skåne and adjacent areas during much of the jurassic (hallam 1994; manspeizer 1994), and vegetation largely covered coastal plains, which resulted in the formation of autochthonous coal beds. these conditions also favoured low-ph chemical weathering in the hinterlands, along sediment transport paths and during pedogenesis. therefore, jurassic strata in skåne include mineralogically mature sandstones, and clay mineral suites with a considerable content of kaolinite. in the late jurassic, the region was subjected to increasing aridity which caused very limited evaporite mineral formation in skåne (caliche nodules). lithostratigraphy the lithostratigraphic scheme of the jurassic succession in skåne (fig. 2) was established by norling (1972, 1981), sivhed (1984) and norling et al. (1993). the jurassic successions are subdivided into the höganäs, rya and annero formations and the informal mariedal formation (fig. 2; norling et al. 1993). exceptions to this do occur. in central skåne, the rhaetian–hettangian is referred to the höör sandstone, which in places is overlain by lithostratigraphically unspecified sediments. in the ängelholm trough, the middle jurassic is referred to the vilhelmsfält formation; in the fyledalen fault zone, the lower jurassic is referred to the informal 529 530 röddinge formation. in the subsurface of south-west skåne, an incomplete succession of jurassic strata ranging from the aalenian to the tithonian has been recorded. this resembles exposed penecontemporaneous strata fairly well, and it is thought likely that parts of the höganäs, rya and annero formations are represented, at least in part, in south-west skåne (norling 1981; d. guy-ohlson, personal communication 1994). höganäs formation stratigraphy and distribution the höganäs formation constitutes a transitional unit between the continental kågeröd formation and the shallow marine rya formation. the formation is wellknown and well-defined in the helsingborg area (sivhed 1984; figs 1, 2). it can be traced in deep wells in southwest skåne (brotzen 1950; larsson et al. 1994) and it is also represented on the skurup platform (thomas et al. 1993). the plant-bearing sediments at rödalsberg and munka tågarp in the fyledalen fault zone (möller & halle 1913) are here tentatively referred to the höganäs formation. it is possible that the members of the höör sandstone in central skåne can be correlated to similar strata of the höganäs formation (troedsson 1951). the höganäs formation is up to 250 m thick in the helsingborg area (sivhed & wikman 1986), 150 m in the ängelholm area (bölau 1959), 40 m in the svedala area (larsson et al. 1994) and probably more than 150 m in the höllviken area. palynomorph dating of the höganäs formation indicates a rhaetian–hettangian age (lund 1977; guy-ohlson 1981). in the ängelholm trough and in the helsingborg area, the höganäs formation is subdivided into the vallåkra and bjuv members (rhaetian) and the helsingborg member (hettangian). landskrona (l) – kävlinge (k)ängelholm trough (ä) helsingborg area (h) sw skåne (s) central skåne (c) vomb trough & fyledalen fault zone (v & f) tithonian kimmeridgian oxfordian callovian bathonian bajocian aalenian toarcian pliensbachian sinemurian hettangian rhaetian norian m id dl e la te ea rl y ju ra ss ic c re t. t ri as si c annero fm annero fm annero fm vilhelmsfält fm mariedal fm mariedal fm rya fm pyroclastic sediments höör sandstonehöganäs fm kågeröd fm ? ? ?in co m pl et e su cc es si on , n ot d ef in ed röddinge fm berriasian fig. 2. lithostratigraphic scheme of the jurassic of skåne, modified from norling et al. (1993). the localities are indicated on fig. 1. 531 sedimentology and petrography the vallåkra member has been reported to comprise poorly stratified kaolinitic and smectitic mudstones and sandstones with sphaerosiderite concretions; the upper parts of the vallåkra member have yielded a restricted marine fauna (troedsson 1948, 1951). the bjuv member is bounded below and above by extensive autochthonous coals. coals, kaolinitic mudstones, heteroliths (flaser and lenticular bedded sandstones and mudstones) and mineralogically mature arenites are common in this member (ahlberg 1994; arndorff 1994). the bjuv member is dominated by floodplain deposits, including mature palaeosols (stacked underclays and autochthonous coals), interbedded with isolated fluvial channel and crevasse splay sandstones (ahlberg 1994; arndorff 1994). in these sediments, numerous dinosaur footprints have been observed (bölau 1952; gierlinski & ahlberg 1993). in addition, sparse evidence of marine influence on the floodplains has been observed, indicating a coastal plain setting. the marine signatures include heteroliths with rhizocorallium isp. and diplocraterion isp. burrows (ahlberg 1994) and dinoflagellates (s. lindström, personal communication 1996). the höganäs formation has traditionally been referred to as deltaic, including 12 deltaic cycles. nine of these are found in the uppermost member, the helsingborg member (troedsson 1950, 1951). this stratigraphic unit includes floodplain strata similar to those of the bjuv member. here, however, the floodplains were subjected to several restricted marine incursions (vossmerbäumer 1969, 1970; ahlberg 1990, 1994; pieńkowski 1991a, b). these are testified by the presence of molluscs, foraminifera (very rare), a well-developed partly marine trace fossil assemblage (fig. 3a), and indications of tidal activity. the tidal signatures include abundant bioturbated muddy heteroliths, which locally show bimodal fig. 3. features of the höganäs (a), rya (b, c) and mariedal (d) formations. a: two diplocraterion isp. burrows associated with rootlets in a muddy heterolith (helsingborg member, helsingborg railway tunnel), representing mixed marine and non-marine bioturbation in a mudflat. b: herringbone structures in marine sandstone (döshult member, laröd road section). c: backscatter electron image of berthierine ooids in thin section (rydebäck member, drill core rydebäck–fortuna 4, at a depth of 67.20 m). d: overturned autochthonous coal bed with pyrite concretions, overlying pedogenically altered sand with rootlets. the parent material (i.e. the light sand without pedogenic overprint to the right) includes diplocraterion isp. burrows. the section probably records infilling of an interdistributary bay and subsequent stabilisation by peat. photo from the top of the fuglunda member, eriksdal sand quarry. a b c d 532 ripple orientation and scoured sand-filled channels. in addition, cosets of cross-bedded sandstones with rhythmically spaced mud drapes and reactivation surfaces have been observed, which indicate uniform westwards (basinwards) sediment transport, possibly in tidallyinfluenced delta distributary channels (fig. 4). the overall facies architecture of the höganäs formation in north-west skåne is dominated by floodplain or muddy bay deposits, which encase laterally restricted sandstone bodies (troedsson 1951; ahlberg 1994). this is suggestive of a high and constant rate of subsidence and vertical sediment accretion (and availability of accommodation space). tabular basinwide sandstone bodies have been described at two levels within the helsingborg member (troedsson 1951). one of these occurs at the base of the hettangian succession, where rhaetian overbank sediments are abruptly followed by hettangian gravel and sand deposited in braided rivers. this change may have been caused by hinterland uplift (troedsson 1951), but could also reflect the upstream effects of a relative sea-level fall in the danish basin. a maximum flooding surface has been recognised in the upper rhaetian of the danish basin, overlain by a sequence boundary and a lower hettangian lowstand systems tract. this low sea-level stand induced basinwards progradation of lowstand shoreface sand, from the margins towards the depocentres of the danish basin (nielsen 1994; nielsen 2003, this volume). such a lowering in stratigraphic base level may have increased the alluvial slope, and triggered the early hettangian onset of braided stream deposition recorded in north-west skåne (ahlberg 1994; see posamentier et al. 1992; shanley & mccabe 1994). in addition, an abrupt increase in surface runoff could have caused a similar depositional pattern (ahlberg 1994). höör sandstone stratigraphy and distribution the höör sandstone crops out in the central part of skåne (fig. 1), where it forms erosional remnants directly overlying weathered precambrian crystalline basement. the 50 m thick formation is of rhaetian–hettangian age, according to its macroand microflora (antevs 1919; troedsson 1940; lund 1977). the threefold division includes a lower unnamed unit, up to 15 m thick, overlain by the stanstorp member (up to 15 m thick; local term: ‘kvarnstenen’ = millstone). the uppermost member, the vittseröd member (up to 25 m thick; local term: ‘byggnadsstenen’ = building stone), is exposed in several abandoned quarries south-east of höör (norling et al. 1993; wikman & sivhed 1993). sedimentology and petrography the höör sandstone outcrops are only a few m2 in size, thus precluding detailed interpretation of depositional setting. the basal unit is presently not exposed, but has been reported to comprise fine-grained sandstones, mudstones and clays (wikman & sivhed 1993). fig. 4. muddy heteroliths and sandy foreset beds with rhythmically spaced mud drapes (arrow) in the helsingborg member. these deposits probably represent bedforms that migrated into an intertidal interdistributary bay. pauses in foreset migration and subsequent mud draping occurred during tidal slackwater (sellwood 1972; terwindt 1981; tyge 1990). section exposed temporarily in 1928 in the centre of helsingborg; photo from the collection of g. troedsson. the middle unit, the stanstorp member, is composed of arkoses, subarkoses and quartz arenites. polymict conglomerates, mudstones and thin coal laminae occur in the unit, and impressions of fossil wood are abundant (troedsson 1940). sedimentary structures include large-scale cross-bedding, cut and fill structures, and a thin, graded event layer with an erosional basal conglomerate and superimposed current ripples, indicating waning energy conditions. a tentative interpretation of the sediments of the stanstorp member is that they were deposited in a continental environment subjected to floods (g. pieńkowski, personal communication 1988). the uppermost unit, the vittseröd member, consists of fine-grained, quartz-cemented quartz arenites and subarkoses. texturally, the sandstones are fairly mature and they have probably been thoroughly reworked by waves and currents. sedimentary structures include trough cross-bedding, which in places form herringbone structures, and rip-up mud clasts on bedding surfaces. plant fragments are not as common as in the stanstorp member, and beds rich in bivalve impressions have been reported (troedsson 1940). the member probably consists of nearshore marine sediments, surrounding islands of weathered precambrian gneiss, forming an archipelago (norling et al. 1993). rya formation stratigraphy and distribution the marine rya formation (fig. 2) is subdivided, from base to top, into the döshult, pankarp, katslösa and rydebäck members. the formation is recognised in the ängelholm, helsingborg, landskrona and kävlinge areas. in south-west skåne, the rya formation is missing or only poorly developed (norling & skoglund 1977; norling 1981, 1982; larsson et al. 1994). based on foraminifers, ammonites and ostracodes, the döshult member is dated to the early sinemurian, the pankarp member to the late sinemurian, the katslösa member to the late sinemurian – early pliensbachian and the rydebäck member to the late pliensbachian – late aalenian (reyment 1969; norling 1972; sivhed 1980, 1984). sedimentology and petrography the döshult member is characterised by coarse-grained sandstones and siltstones in the lower part, and is dominated by clays and marls rich in marine fossils in the upper part. the member is up to 80 m thick in the ängelholm, helsingborg and landskrona areas (bergström et al. 1982). at present, the basal part of this member is exposed at three localities in the helsingborg area (for locations, see norling et al. 1993). these contain mineralogically and texturally mature, trough crossbedded sandstones, commonly with herringbone structures showing north and south oriented palaeocurrent directions (fig. 3b). the occurrence of herringbone structures in well-sorted sand (tidal or not) suggests high energy foreshore to subtidal marine depositional conditions for the lower part of the member. in an abandoned quarry in north-west skåne (gantofta brickpit in the helsingborg area, outcrop very limited at present) the upper part of the döshult member commences with bioturbated marine nearshore sands, including diplocraterion isp., rhizocorallium isp., chondrites isp. and planolites isp. burrows, as well as abundant marine invertebrate body fossils (frandsen & surlyk 2003, this volume). this is followed by a bioturbated shelf mudstone with storm-deposited sand and silt intercalations (tempestites). a massive red mudstone with scarce marine body fossils and burrows follows, which is interpreted as having been deposited rapidly, in a low energy but oxidising environment. the youngest part of the succession comprises siltstones and mudstones, with carbonate-rich beds, deposited in a shallow marine setting (rolle et al. 1979). the pankarp member has an estimated thickness of up to 70 m in the subsurface of the ängelholm, helsingborg and landskrona areas. in the kävlinge area, the thickness is about 20 m (sivhed 1980; norling 1981; norling et al. 1993). in westernmost skåne, the pankarp member has been observed in small diameter drill cores. there, the member is subdivided into a lower unit of variegated clays and shales, a middle, poorly sorted silty to sandy unit including a coal bed, and an upper monotonous mudstone unit which is silty and organicrich at the base, and reddish–greenish at the top (sivhed 1980; norling et al. 1993). in the uppermost part of one core, the pankarp member comprises lenticular bedded heteroliths with planolites isp. burrows. the katslösa member is mainly known from the subsurface in westernmost skåne, and it has a thickness of 30–40 m in the ängelholm, helsingborg and landskrona areas. in the kävlinge area, the thickness is about 75 m (sivhed 1980; norling 1981; norling et al. 1993). the most complete section, at katslösa in the helsingborg area, was described and sampled by troedsson (1951), as he followed a temporary trench dug perpendicular to the strike of the tilted strata. sedimento533 logical interpretations given here are mainly based on the results of recent petrographical studies of museum collections, combined with published lithological descriptions (troedsson 1951; norling 1972). the katslösa member yields a rich marine microfauna and macrofauna, and it is dominated by homogeneous mudstone deposited in a marine low-energy environment. thin beds of matrix-rich quartz wackes are common. they are typically mineralogically mature but texturally highly immature with abundant angular sand grains. the matrix comprises organic matter, micrite, mica and clay minerals. in thin section, the sandstones show evidence of intense burrowing, which has obliterated depositional structures. scattered berthierine ooids, as well as authigenic siderite crystals have been observed. the rydebäck member is up to 70 m thick in the ängelholm, helsingborg and landskrona areas. it is only known from subsurface material in westernmost skåne, and sedimentological conclusions herein are entirely based on observations from two wells (drill cores rydebäck–fortuna-1 and -4). the member comprises a uniform succession of muddy arenites with a rich marine microfauna (mostly foraminifera), and represents deposition in an offshore low-energy environment (norling 1972). the sediments are strongly burrowed, which has caused an effective mixing of sand and mud, resulting in the forming of quartz wackes. the sand is quartz-rich, and grains are typically well rounded. berthierine ooids are common constituents of the sediment (fig. 3c). in conclusion, deposition of the rya formation began with nearshore coarse clastics, and continued with offshore mudstones with tempestites (the döshult member), followed by offshore muddy sediments with a brief nonmarine interval (the pankarp member), and ended with deposition of open marine low-energy deposits (the katslösa member and the rydebäck member). hence, the marine rya formation shows an overall fining-upwards trend, and an up-section bathymetric deepening of the depositional environment. it is notable that the stormdominated, hummocky cross-stratified hasle formation on bornholm is contemporaneous with the muddy katslösa member of the rya formation (surlyk & noenygaard 1986; koppelhus 1995). the depositional environment in western skåne was either physically protected from the storm energy due to basin topography, or deposition in skåne took place below storm wavebase. berthierine ooids occur scattered in the katslösa member and are increasingly abundant up-section in the rydebäck member. there is an intriguing possibility that iron ooid formation was promoted by precipitation of iron and silica from volcanic exhalative fluids rising up through the substrate, as has been reported from modern marine sediments offshore indonesia (heikoop et al. 1996). this hypothesis has emerged with the recent publication of new age data for the volcanic rocks in skåne, which now appear to be comparable in age to the prominent iron ooid-bearing deposits, i.e. the rydebäck member and the röddinge formation (see below). röddinge formation stratigraphy and distribution due to the limited degree of exposure and the weathered state of the surficial strata, the informal stratigraphic unit ‘röddinge formation’ was introduced for the iron-rich deposits in the fyledalen fault zone, primarily at kurremölla (norling et al. 1993). the röddinge formation is largely known from museum collections, and it has been dated by ammonite biostratigraphy to be of sinemurian–pliensbachian age (moberg 1888; reyment 1959). based on unpublished museum data (sample locations, etc.), it is calculated that the röddinge formation has a thickness of several hundreds of metres. petrography the röddinge formation is dominated by fineto medium-grained quartz arenites, with subordinate thin conglomerates. the sediments are moderately consolidated by berthierine or siderite cement, and berthierine oolites are common in the succession. the ooids are generally around 0.3 mm in diameter and ellipsoidal in shape. the core commonly consists of detrital quartz or heavy minerals, surrounded by concentric microlaminae of berthierine. the deposits are strongly affected by modern weathering and are characterised by a red, brown or yellow stain (iron hydroxides). unweathered deposits, known from boreholes, are greyish dark green due to the content of berthierine and siderite. the iron content is 8–10% in the weathered sandstones, up to 20% in the oolites, and at kurremölla a 1.7 m thick oolite bed has an iron content of up to 35%. this ore bed was mined during the 1930s (hadding 1933). due to the poor, weathered exposure, the röddinge formation has not been subjected to detailed facies 534 analysis. in general, the mineralogical and textural maturity of the sands imply prolonged reworking. marine influence is clearly indicated by finds of ammonites and crinoids (hadding 1933). the shape of the iron ooids indicate diagenetic precipitation, prior to and during sediment compaction. as discussed above with respect to the berthierine ooids in the rydebäck and katslösa members, volcanic activity may have stimulated the process. lateritisation has been suggested as another possible source for the iron enrichment in the sediments (nilsson 1992). vilhelmsfält formation and mariedal formation stratigraphy and distribution middle jurassic sediments are found in parts of western skåne, in the fyledalen fault zone and in the ängelholm trough. the deposits are divided into two formations, the vilhelmsfält formation and the mariedal formation (norling et al. 1993). the vilhelmsfält formation is restricted to the ängelholm trough. it has been dated as bajocian–bathonian on palynological grounds (guy-ohlson 1971, 1986, 1989). similarities between the vilhelmsfält formation and the fuglunda member of the mariedal formation were noted by bölau (1959), but such a correlation cannot be confirmed at present. use of the informal name ‘mariedal formation’ was recommended by norling et al. (1993), until the identity of the formation is resolved. the mariedal formation crops out in the fyledalen fault zone, and is also found in the landskrona and helsingborg areas (fig. 1). it has been divided into the fuglunda member and the overlying glass sand member, which are dated to the bajocian and the bathonian, respectively, on palynological and palaeobotanical grounds (tralau 1966, 1968). the fuglunda member is up to 75 m thick, whereas the glass sand member has a thickness of up to 100 m (norling et al. 1993). lithology of the vilhelmsfält formation the vilhelmsfält formation consists mainly of sandy, micaceous mudstones with plant remains and sandy intercalations. a thin coal bed has been observed in a sandy layer in the lower part of the formation. the formation is more than 75 m thick (bölau 1959). lithology and sedimentology of the mariedal formation in a quartz sand quarry at eriksdal in the fyledalen fault zone, the fuglunda member comprises thirteen deltaic cycles of coals, heteroliths and sandstones (koch 1979; rolle et al. 1979). the coals are commonly associated with rootlet horizons and immature palaeosols (fig. 3d). hence, they were formed autochthonously, primarily by in-situ accumulation of wood in peat swamps. indeed, between each coal bed, several soil formation ‘attempts’ (i.e. immature palaeosols) may be distinguished. the palaeosols differ clearly from their parent material in terms of colour, texture and degree of stratification. typically, the lower boundaries are gradational, and the upper boundaries are sharp. rooting and eluviation of clay minerals and organic matter are additional pedogenic characteristics (allen & wright 1989). the heteroliths include the full spectrum from almost pure mudstone (rich in clay-sized coal detritus) with slightly lenticular silt laminae, through lenticular, wavy and flaser bedded heteroliths. a partly marine trace fossil assemblage is characteristic of the heteroliths, including diplocraterion isp. and monocraterion isp. the bioturbated muddy heteroliths interfinger with sandy heteroliths, autochthonous coals and palaeosols. they were probably deposited in intertidal low energy environments, i.e. interdistributary bays (rolle et al. 1979). abundant large (< 10 cm) pyrite concretions occur in the coals and the organic-rich heteroliths, implying a marine influence on the eodiagenesis (curtis & coleman 1986). flaser bedded, sandy heteroliths commonly show trough cross-bedding, mud drapes, climbing ripples and channel scour. this facies probably represents deposition in tidal channels. in the eriksdal quarry, a coarse-grained, pebbly sandstone showing scour-andfill structures, forms a lenticular body encased in the coals and heteroliths of the fuglunda member. this facies probably represents a delta distributary channel fill. in conclusion, the fuglunda member at eriksdal shows typical features of a delta plain environment. the glass sand member (up to 100 m thick) succeeds the fuglunda member in the eriksdal quartz sand quarry. the sediments consist predominantly of coarseand fine-grained sandstone (or sand), with minor intercalations of heteroliths and mudstone with diplocraterion isp. burrows (rolle et al. 1979). differing interpretations have been offered for the depositional setting of these sediments. the member was described by rolle et al. (1979) as being of foreshore and lagoonal origin. 535 536 clay silt sand clay silt sand clay silt sand clay silt sand 1 m c b a c d d d fig. 5. log from the glass sand member at eriksdal (j. johansson and l. adrielsson, personal communication 1998). a, shallow stream or unconfined flow (sheet flood) deposits from crevasse splays close to an active channel; b, distal crevasse splay deposits in extremely shallow interdistributary bay (or floodplain) area; c, turbiditic deposition in standing water, probably by subaqueous overbank floods; d, distributary channel deposits showing two-dimensional dunes with superimposed ripples, three-dimensional dunes and cross-channel bars implying unidirectional flow; e, channel erosion and lateral migration with slump sediments (massive sand and large peat (coal) balls, i.e. reworked floodplain swamp deposits) from river bank collapse; f, channel fill deposits dominated by ripple-laminated sand, fining-upwards; g, channel fill concluded and establishment of flora. f. surlyk (personal communication 1998) considers the glass sand member to represent high-energy tidallydominated environments ranging from tidal inlets and open estaurine conditions with large tidal sand bars, flats and channels to marine foreshore; the upper boundary of the member is interpreted to record a major relative sea-level fall. in contrast, j. johansson and l. adrielsson (personal communication 1998) investigated a section through the glass sand member and concluded that fluvial deposits prevail throughout, with few signs (e.g. trace fossils) indicating proximity to the marine realm (fig. 5). these workers thus proposed that the succession probably represents the fluvial-dominated part of a delta plain. 537 clay silt sand clay silt sand clay silt sand massive bedding, sandstone horizontal lamination normal grading planar cross-stratification trough cross-stratification scour-and-fill structures ripple cross-lamination, trough ripple cross-lamination, tabular mudstone reworked peat ball slumping trace fossils (diplocraterion) roots d f e e d f g annero formation stratigraphy and distribution from below, the annero formation (fig. 2) is subdivided into the fortuna marl, the fyledalen clay, the nytorp sand, and the vitabäck clay. based on biostratigraphy (foraminifera, ostracodes and palynomorphs), the fortuna marl is referred to the bathonian–oxfordian, the fyledalen clay to the oxfordian–kimmeridgian, the nytorp sand to the kimmeridgian–tithonian, and the vitabäck clays to the tithonian–berriasian (christensen 1968; norling 1972, 1981; erlström et al. 1991). sediments referred to the annero formation are found in western skåne including the ängelholm trough, in the fyledalen fault zone and in the vomb trough, and have been described by norling (1972, 1981) and erlström et al. (1991). the fortuna marl is restricted to the landskrona– kävlinge area, where it has a thickness of up to 10 m (guy-ohlson & norling 1988). the fyledal clay is about 30 m thick in the helsingborg–landskrona area and in the ängelholm trough, and more than 55 m thick (estimated to 140 m) in the fyledalen section, and 24 m thick in the vomb trough. the nytorp sand is 20–25 m thick in the helsingborg–landskrona area, 20 m thick in the fyledalen section and 5 m thick in the assmåsa borehole. the vitabäck clay is 20–25 m thick in the ängelholm and vomb troughs, more than 5 m thick in the landskrona–helsingborg area, and more than 34 m thick in the fyledalen section (norling 1981). lithology and petrography the thin fortuna marl has only been observed in a few drill cores of limited quality, and sedimentological information is therefore minimal. the strata comprise sandstone, carbonate-rich siltstones and mudstones, with a marine microfauna (guy-ohlson & norling 1988). the fyledal clay is characterised by a uniform succession of greenish and organic-rich clays, interrupted by thin silt and sand beds (guy-ohlson & norling 1988; erlström et al. 1991). the sediment displays, on the one hand, continental and freshwater indicators such as rootlet beds, oligohaline–mesohaline ostracodes and caliche nodules. on the other hand, marine indicators are also recorded, such as calcareous foraminifera. erlström et al. (1991) thus suggested that deposition mainly took place on a muddy coastal plain, with stagnant ponds (lagoons). fining-upwards sandstone wedges (< 1 m thick) with erosional bases have been observed in the fyledal clay at the eriksdal quartz sand quarry. these include gravel with rip-up mud clasts and mollusc shells at the base, as well as reworked glauconite in their higher parts. the sand wedges have been interpreted as washover fan deposits (erlström et al. 1991). the occurrence of upper jurassic evaporites and euryhaline biota are very few, but may be of particular significance. whereas the caliche nodules are well preserved, it is unclear if the gypsum is of primary origin. in any case, these represent the only signs of aridity in the jurassic successions of skåne. hence, skåne may have been only marginally influenced by the late jurassic north-eastwards migration of the arid climate zone of western and central pangea (hallam 1994). the nytorp sand is not accessible for facies analysis. however, due to its stratigraphic position between the fyledal clay and the similar vitabäck clay, erlström et al. (1991) inferred deposition in a back-barrier environment. the vitabäck clay, which straddles the jurassic– cretaceous boundary, is primarily known from lithological and palaeontological investigations of samples, as continuous sections are not available. the mud-dominated formation shows similarities to the fyledal clay, only with a slightly higher content of sandy beds. palynomorphs, ostracodes and molluscs indicate nonmarine conditions with brackish marine influence, and deposition probably took place on a muddy coastal plain (ekström 1985; erlström et al. 1991). jurassic volcanism and pyroclastic deposits in central skåne, numerous mesozoic basaltic necks occur (norin 1933). the basalts are typically vitrophyric with columnar structures, and undersaturated with respect to silica (see wikman & sivhed 1993 and wikman et al. 1993 for reviews). k-ar datings of the basalts indicate bathonian and aptian ages (printzlau & larsen 1972; klingspor 1976). recent palaeomagnetic datings, however, have yielded older, toarcian–aalenian ages (bylund & halvorsen 1993). despite the abundant volcanic necks in skåne, the distribution of pyroclastic deposits appears surprisingly limited. this, and the abundance of lithified magma spray droplets, may indicate a hawaiian or strombolian type of eruptive activity, i.e. rather passive eruptions of low-viscosity basaltic magma (scarth 1994). only locally, around the volcanic necks of central skåne, up to 50 m thick volcanic tuffites occur, which presently are exposed 538 in two sections (norling et al. 1993). no correlative pyroclastic beds have yet been identified in sedimentary basins surrounding central skåne. the tuffites are dominated by lapilli, which in thin section show clusters of tiny spherical condensed glass particles which are now diagenetically transformed to clay minerals and other relatively stable minerals. other constituents comprise rare volcanic bombs (typically 5–10 cm in cross-section), angular and rounded gneiss boulders (typically 5–15 cm in cross-section), probably from the sidewall, scattered detrital quartz grains, and pieces of fossil wood, including large logs. palynological dating of the tuffites has yielded an age at the early–middle jurassic boundary (tralau 1973). in central skåne, the volcanic intrusions have probably had a profound effect on the diagenesis of the lower jurassic arenites (see below). diagenesis, thermal history and petroleum potential in the jurassic strata of skåne, most diagenetic properties were obtained rather soon after deposition, i.e. eodiagenetically. hence, the sediments were influenced by surficial features such as climate, groundwater and pedogenic processes. due to only moderate burial temperatures, burial diagenetic features are less pronounced (ahlberg 1996). a low thermal impact has been confirmed by guy-ohlson (1989; personal communication 1994) who observed low thermal alteration index (tai) values of palynomorphs (around 3) throughout the jurassic of skåne. the rhaetian–hettangian höör sandstone in central skåne is an exception to this. as revealed by cathodoluminescence microscopy and fluid inclusion analysis, the höör sandstone was subjected to flushing of hot hypersaline brines, which probably were released in association with the early–middle jurassic volcanic events (ahlberg 1994). the organic content of the jurassic strata in skåne is typically dominated by gas prone kerogen (type iii), which is below, or at the onset of, thermal maturity (ahlberg 1996). as the burial heat was insufficient, hydrocarbon maturity can only have been achieved locally, possibly in association with faulting and volcanic activity. the faults, however, were also potential escape routes for migrating hydrocarbons (ahlberg 1996; ahlberg & olsson 2001). references ahlberg, a. 1990: provenance, stratigraphy, palaeoenvironments and diagenesis of the lower jurassic strata in the helsingborg railway tunnel, southern sweden, 2, 54 pp. unpublished licentiate thesis, lund university, sweden. ahlberg, a. 1994: deposition and diagenesis of the rhaetian– hettangian succession (triassic–jurrassic) in southern sweden: a synthesis. lund publications in geology 123, 16 pp. ahlberg, a. 1996: petroleum in the mesozoic of sweden – why not? 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geologiske undersøgelse serie a 34, 23 pp. vossmerbäumer, h. 1969: paläoökologische ausdeutung fossiler wurzelböden. geologiska föreningens i stockholm förhandlingar 91, 112–126. vossmerbäumer, h. 1970: untersuchungen zur bildungsgeschichte des unteren lias in schonen (schweden). geologica et palaeontologica 4, 167–193. wikman, h. & sivhed, u. 1993: beskrivning till berggrundskartan kristianstad sv. sveriges geologiska undersökning serie af 155, 106 pp. wikman, h., bergström, j. & sivhed, u. 1993: beskrivning till berggrundskartan helsingborg so. sveriges geologiska undersökning serie af 180, 114 pp. ziegler, p.a. 1990: geological atlas of western and central europe, 2nd edition, 239 pp. amsterdam: elsevier for shell internationale petroleum maatschappij. manuscript received 2 september 1994; revision accepted 31 may 2000. e2019430204-01 basement highs are large structural features, commonly buried in sedimentary basins (busby & azor 2012). they are of interest for natural resources exploration and research because of their ability to influence migration and entrapment of petroleum (trice 2014) and water, and the deposition of metals (hitzman 2005; borg et al. 2012). three-dimensional (3d) reservoir models (e.g. shepherd 2009) are built to evaluate and model fluid-filled basement reservoirs (ringrose & bentley 2015). however, subsurface data are expensive, difficult to obtain and are often widely spaced. ideally, basement reservoir models would be constrained by rock, fracture and mineral vein data from appropriate outcrop analogues (acknowledging that subaerial basement rocks have, by definition, a different uplift history than subsurface basement). the liverpool land basement high (llbh) in greenland is an uplifted and well-exposed basement high located between two sedimentary basins, and thus provides a valuable analogue for fractured basement-hosted mineral, oil and geothermal reservoirs. the geological survey of denmark and greenland (geus) conducted reconnaissance work on the llbh in 2018 to assess the quality of the exposure of basement palaeo-weathering profiles and fault-fracture networks. here, we introduce the llbh, the concept of fractured basement liverpool land basement high, greenland: visualising inputs for fractured crystalline basement reservoir models graham banks*1, stefan bernstein1, sara salehi1, pierpaolo guarnieri1, dennis bird2, catherine hamblett3, david peacock4 and jon foster5 geus bulletin is free to individuals and institutions in electronic form. the author(s) retain copyright over the article contents. review article | open access geus bulletin vol 43 | e2019430204 | published online: 22 july 2019 https://doi.org/10.34194/geusb-201943-02-04 viking graben faroe-shetland basin møre basin vøring basin m ør etr øn de la g fa ul t c om pl ex m t g g f h bf su f re yk ja ne s ri dg e k im be rly r id ge ae gi r r id ge w jan mayen fz e jan mayen fz jan mayen microcontinent tjörnes fz systemgreenlandiceland ridge icelandfaroe ridgeiceland shelf iceland basin 60°n 10°e pzeg rockall basin ? ? oceanic elements: high (including platforms, terraces and intra-basinal elevations) basin continental elements: continent-ocean boundary active mid-ocean ridge rift system normal fault reverse fault strike-slip/transcurrent fault fault (unspecified) 1:10,000,000 0 100 200 300 km geodetic reference system: wgs 1984 projection: lambert conformal conic with central meridian: -40, std. parallels: 55, 75 uk iceland g re en la nd n or w ay lf uh fig. 2b fig. 2a llbh llb n jlb fig. 1. locations of north atlantic ocean sedimentary basins (blue) and highs (pink). liverpool land basement high (llbh), liverpool land basin (llb) and jameson land basin (jlb) are located on the western north atlantic conjugate margin. the lancaster field (lf) and utsira high (uh) basement petroleum reservoirs are located on the eastern north atlantic conjugate margin (modified from hopper et al. 2014). https://doi.org/10.34194/geusb-201943-02-04 e2019430204-02 reservoir modelling, and how studying the llbh can help enhance reservoir modelling of fractured basement. we present some of our preliminary observations of llbh faultfracture networks and discuss how the exposed sedimentbasement features and processes might aid industry and research in their top basement mapping activities. we propose that llbh provides a particularly suitable analogue for industry and research to analyse: (a) multiscale fracture system connectivity, (b) fluid migration and fluid-rock reaction processes, (c) input parameters for basement reservoir modelling and (d) top basement geomorphologies and processes. ‘basement’ is here defined as any crystalline lithology and associated regolith. for definitions of fracture, fracture corridor, fault, fault zone, joint, vein and aperture see peacock et al. (2016). liverpool land basement high llbh is a precambrian–caledonian (corfu & hartz 2011) crystalline basement massif onshore central east greenland (71on, 22ow), located on the western north atlantic margin, between the offshore liverpool land basin (llb) and the onshore jameson land basin (jlb; figs 1, 2). it is mostly comprised of granites, monzodiorites, gneisses and schists, and is the result of multiple tectonic, metamorphic, magmatic, burial, sedimentation, uplift and erosion events (corfu & hartz 2011; guarnieri et al. 2017). glacial erosion has generated a deeply incised outcrop 125 × 35 km, rising up to 1.2 km a.s.l. llbh has almost no vegetation cover, so rock exposures continuously display 3d relationships between lithology, structure, weathering and basin sedimentation across a range of scales, from tens of kilometres, down to millimetres (fig. 3). llbh and neighbouring jlb have been explored for decades. sedimentary copper showings discovered in permian–triassic strata by nordisk mineselskab and avannaa resources–anglo american were drilled by the latter consortium in 2014 (guarnieri et al. 2017). jlb hydrocarbon seeps attracted petroleum exploration by arco (atlantic richfield company; guarnieri et al. 2017) and most recently by greenland gas & oil (greenland gas & oil 2019). much tectonic, lithology and petrology research has been conducted upon and offshore llbh. however, to our knowledge, this review is the first to highlight the potential of the llbh exposure to visualise input parameters for sea level 0 4 8 12 16 plio-pleistocene neogene paleogene cretaceous permian–jurassic oceanic crust devonian – m. permian paleogene sills 25 km nw se jlb llbh llb km pre-permian w e twt 1000 2000 viking graben haugaland high luno graben augvald graben avaldnes high 3000 edvard grieg discovery well johan sverdrup discovery wellrolvsnes well (projected) palaeozoic early cretaceous basement johan sverdrup field late cretaceous edvard greig field jurrasic/triassic paleocene owcowc 5 10 km2.50 fig. 2. a: schematic geoseismic section (depth) across jameson land basin (jlb), liverpool land basement high (llbh) and liverpool land basin (llb) on the western north atlantic margin (modified after hamann, 2005). fig. 2. b: interpreted seismic section (in two-way time) across the southern utsira high basement oil fields, eastern north atlantic margin. the position of the rolvsnes petroleum discovery in fractured basement is indicated in b. see fig. 1 for section locations. (reproduced with permission from lie et al. 2016). e2019430204-03 fractured basement modelling of petroleum, geothermal and mineral commodities. fracture systems in crystalline basement reservoirs natural fracture systems are crucial components of basement reservoir models. fracture systems are ubiquitous in upper crustal rocks and include faults, joints, veins and fracture apertures. they dominate the hydrological behaviour of rocks (e.g. fig. 4 in belaidi et al. 2018) – especially in lithologies with low matrix porosity where fluid flow is channelised through permeable fractures (vidal et al. 2017). fracture systems enable basement highs to transmit or trap petroleum – the lancaster field, offshore uk (belaidi et al. 2018; fig. 1) and the utsira high, offshore norway (riber et al. 2015; fig. 1) are two examples. likewise, geothermal energy is often extracted from ‘hot’ water in fractured basement (e.g. vosgerau et al. 2016; vidal et al. 2017). mineral deposits that precipitated into basement fracture networks during hydrothermal mineralisation occur worldwide (sheldon & micklethwaite 2007; walter et al. 2018). therefore, modelling natural fracture systems is fundamental to predicting reservoir capacity, petroleum, groundwater and geothermal resources, and to estimate the amount of vein-hosted mineral resources. reservoir modelling of fractured basement: data resolution challenges fractured reservoirs are complex, multi-component entities. it is often difficult to characterise the interactions of all their components in the subsurface and to quantify how fluids might move through them. therefore, 3d reservoir models are built to quantitatively display and integrate subsurface parameters (shepherd 2009; ringrose & bentley 2015). applications include estimating trapped fluid volumes, simulating fluid flow and planning where to drill wells. when building a 3d model of a fractured basement reservoir, large-scale basement faults can be interpreted on 3d seismic data (fig. 4), but most lithology and fault-fracture network properties, e.g. spacing, apertures, porosity, permeability and fluid saturation are below seismic resolution. characterising sub-metre-scale lithologies and fractures is conducted at the wellbore (e.g. with borehole images, core plug and sonic data), but upscaling and interpolating a fracture system’s properties away from the wellbore would be model-driven and highly uncertain. moreover, some reservoir modellers may not have actually studied extensive 3d fracture systems in the field. the llbh is a particularly suitable location for fractured basement reservoir geoscientists, modellers and engineers to understand, constrain, extrapolate and interpolate fractured basement properties. 2018 llbh reconnaissance: initial observations remote sensing and field observations made in 2018 documented the suitability of the llbh for characterising fracture networks and weathering features needed for a fractured basement reservoir model. first, a reconnaissance analysis was conducted using landsat-8 oli, sentinel-2 and aster spaceborne data, which enabled rapid identification of regional-scale structural lineaments. fieldwork was then conducted to examine the exposure quality of a range of finer-scale structures (figs 5a, b, c, d), mostly in an area c.10 km east of nerlerit inaat (constable pynt) airport and supplemented with aerial observations around llbh by helicopter. all llbh continuum of rock outcrop scales on the liverpool land basement high seismic imaging colour coding of fractured reservoir data types horizontal resolution and range (m) inter-well electro-magnetic logs dynamic flow data single well logs core limit of most outcrops fluid, rock, & fracture fluid ve rt ica l r es ol ut io n an d ra ng e (m ) 0 1 10 100 1000 10 000 100 000 0 1 10 100 1000fig. 3. spatial scales of input data for fractured reservoir models. geological outcrops are often limited to the scales indicated by the red dashed line. llbh encompasses the full range of scales required to cover all fractured reservoir para meters (purple dashed line). e2019430204-04 lithologies are pervasively fractured (figs 5a–d). for example, hurry inlet granite outcrops host 100 m wide corridors of planar joints and chlorite-lined strike-slip fault systems, whilst the janus ø schist (fig. 5b) has fractures that reach more than 200 m in height. vertical variations in the spacing of fractures in the schist reveal their mechanical stratigraphy. llbh is segmented by generally east–west-oriented glacial valleys and fjords, with some cliff outcrops more than 500 m high. these outcrops display fracture networks at a scale suitable for reservoir model cross-sections. below seismic scale (figs 5b, c, d), the rock exposures continue to display fracture network-scaling relationships (e.g. fracture length-height ratios and the number of short versus long fractures) that are vital to realistically populate a fractured basement reservoir model. the frequencies and widths of fault zones and fracture corridors can be observed in 3d (figs 5a, b). the sparse regolith and vegetation enable systematic analysis of structures, chemical alteration and mineralisation progressively across the fault zones. between the fault zones and fracture corridors, 10 to 100 m size blocks of basement ‘matrix’ are pervasively fractured from metreto millimetrescales regardless of lithology (figs 5c, d). some of the ubiquitous llbh epidote-quartz-carbonate veins are only partly cemented (fig. 5d). such rugose and cement-bridged apertures could maintain effective porosity-permeability in crystalline rocks at reservoir conditions. this shows that whilst some subsurface veins may act as fluid barriers, others could actually contribute to a reservoir’s fluid storage and drainage. the sub-metre joints and partly cemented veins also indicate that llbh is an analogue for dual-porosity basement reservoir types (i.e. both discrete and non-discrete fracture sets contribute to fluid flow; e.g. belaidi et al. 2018). characterising llbh veins can help visualise how to incorporate mineralisation and dissolution history into basement reservoir matrix permeability models. a reservoir’s top basement bedrock and regolith zone may be able to store significant volumes of fluid or mineral resources (holland 2011) and should be studied in detail. subsurface seismic interpretations often involve mapping top basement, but this is frequently schematic due to limited top basement resolution or limited knowledge of top basement characteristics. the well-exposed and incised top basement of llbh (marked by the nonconformity between jlb sedimentary rock and underlying basement) enables detailed study of constituent geomorphologies, weathering facies, regolith types and porosity-permeability networks formed during arid peneplanation (permian) and glacial (quaternary) processes. regolith types include talus slopes, proximal conglomerates, freeze-thaw rubble, proximal outwash plains and kaolinitised peneplains. top basement in the kangerterajittap ilinnera (klitdal) area forms a nonconformity contact between jlb pingodal formation basal conglomerate and gneiss. here, some of the orthogonal joints in the pingodal formation appear to be non-stratabound, i.e. continue into underlying gneisses, and could have connected basin and basement fluid systems. conclusions our 2018 reconnaissance suggests that llbh is a particularly suitable location for petroleum, mineral and geothermal resource sectors to visualise connectivity of dual-porosity basement reservoir fracture networks and mineralisation features continuously across a range of scales. studying the llbh could help geoscientists, modellers and engineers to formulate procedures to upscale wellbore data and downscale seismic data in their fractured basement reservoir models. llbh also provides drilling and reservoir engineers with visual insights into fractured basement heterogeneity. the llbh top basement displays a range of palaeo-weathering y-axis y-axis x-axis x-axis z-axis z-axis fig. 4. fault planes interpreted from seismic data are just one input type for a fractured basement reservoir model. this example is from the lancaster field (see fig. 1. for location). the grid cells are 1 km3. the arrow points north, and the green face points to shallower depths. modelling the intervening reservoir matrix can be facilitated by analysing appropriate analogue outcrops. (reproduced with permission from belaidi et al. 2018). e2019430204-05 profiles and glacial geomorphology features to aid any top basement seismic interpretation. the vein mineralisation and dissolution histories exposed on llbh can help discern the effects of basement highs on basinal brine migration processes during basinand basement-hosted base metal mineralisation. the following projects are being considered for further study: • to collect and interpret photographs of valley/fjord cliffs to 3d-map the laterally continuous fracture networks, and create reservoir-scale virtual outcrops, using photogeological techniques (e.g. sørensen and dueholm 2018). • organise a workshop on the llbh in the field to facilitate discussions between reservoir geoscientists, geomodellers and reservoir engineers of fractured basement reservoir modelling and fracture systems. • investigate the permian-triassic palaeoenvironments of the jlsb-llbh nonconformity to understand top basement geomorphologies, evolution and reservoir properties. acknowledgements we thank the permian-triassic of east and north-east greenland project at geus for sponsorship; the geus logistics and field equipment team; staff at nerlerit inaat (constable pynt) airport and the air greenland helicopter crews for their contributions to the 2018 fieldwork. we also thank tim needham (needham geoscience limited) and tom wesby (first quantum) for their valuable reviews of this manuscript, jan lie ew c. 200 m c. 15 m ew b d c c. 200 m looking east 2 cm a a fig. 5. reconnaissance of the llbh. a: exposure of top basement morphology and 3d fracture network connectivity. b: pervasive fractures > 200 m high on janus ø. c: intense fracturing of basement matrix at the scale of one reservoir model grid cell (sub-kilometre scale). d: partly-cemented vein aperture at core-plug scale (centimetre-scale). e2019430204-06 *corresponding author: graham banks | e-mail: gb@geus.dk 1 geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350, copenhagen k, denmark. 2 department of geological sciences, stanford university, stanford, california 94305, usa. 3 jægersborg alle 55, 2920 charlottenlund, denmark. 4 department of earth science, university of bergen, allégaten 41, 5007 bergen, norway. 5 ineos oil & gas denmark, teknikerbyen 5, 2830 virum, denmark. (lundin-norway) for supplying fig. 2b, daniel bonter and robert trice (hurricane energy) for supplying fig. 4. references belaidi, a., bonter, d.a., slightam, c. & trice r.c. 2018: the lancaster field: progress in opening the uk’s fractured basement play. in: bowman, m. & levell, b. (eds): petroleum geology of nw europe: 50 years of learning – proceedings of the 8th petroleum geology conference, 385–398. borg, g., piestrzynski, a., bachmann, g. h., püttmann, w., walther, s. & fiedler, m. 2012: an overview of the european kupferschiefer deposits. in: hedenquist, j.w., harris, m. & camus, f. (eds): geology and genesis of major copper deposits and districts of the world: a tribute to richard h. sillitoe. special publications – society of economic geologists 16, 455–486. busby, c. & azor, a. 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only minor fine-grained matrix is present. sandstones of lithofacies i have a low k content and most of the k is hosted in feldspar. porosity varies between 23% and 28% and permeability is in the range 200–2000 md. lithofacies ii sandstones have a grain-supported texture, with a predominance of quartz grains; fine-grained matrix fills the intergranular volume. sandstones of lithofacies ii have an intermediate k content, with k-feldspar, mica, and illite as the main sources. porosity varies between 11% and 17% and permeability is in the range 0.4–25 md. lithofacies iii has a matrix-supported texture with quartz grains floating in a clay-rich matrix. samples from lithofacies iii have the highest k signal. illite and illitised kaolinite are roughly equal in importance as sources of k. porosity is up to 11% and permeability up to 0.5 md. the th and u content of all lithofacies is governed primarily by the presence of heavy minerals; no apparent general relationship between u and total organic carbon (toc) was found. comparisons between the core measurements of k, th, and u, and the sng log disclosed a discrepancy between the calibrations of laboratory and borehole measurements. for u the discrepancy contains an erratic element, whereas the difference for k and th can be eliminated by correction factors. thus, the conclusions based on laboratory measurements appear to be applicable to the log data, and, using corrected k values, the facies subdivision can be extended throughout the reservoir section based on the sng log. keywords: danish central graben, north sea, middle jurassic, sandstone reservoir, gamma-ray spectometry, mineralogy, geochemistry i.l.f., l.d.f.*, a.s.‡ & u.k., technical university of denmark (dtu), dk-2800 lyngby, denmark. *present address: internationalt patent-bureau, dk-2630 tåstrup, denmark. ‡present address: jar›frø›isavni›, brekkutún 1, fo-110 tórshavn, faroe islands. e-mail: ilf@er.dtu.dk geological survey of denmark and greenland bulletin 1, 349–366 (2003) © geus, 2003 350 natural gamma-ray wireline logging is commonly used for lithological, or more precisely mineral, identification in boreholes. in sandstone reservoir sections, its main purpose is to obtain an indication of the clay content and thus, reservoir quality. schlumberger (1982) has described two types of gamma-ray logs: the conventional gamma ray log records the total natural gamma radiation in the borehole independent of the energy of the gamma rays, whereas the spectral natural gamma-ray (sng) probe measures the gamma-ray spectrum, i.e. the energy distribution of the gamma rays. the gamma rays originate from the decay of radiogenic potassium (40k), thorium (232th) and uranium (238u), each of which exhibit a characteristic energy spectrum. the measured energy spectrum is transformed into concentrations of k, th and u (schlumberger 1982). in the west lulu-3 appraisal well of the middle jurassic harald field, sng logging was performed in order to differentiate micaceous sand from shale. this distinction should in theory be possible from the th/k ratio, but the picture is complicated because several radioactive minerals are present in the micaceous sandstones that give a more shale-like response. the purpose of this study was to identify the radioactive minerals in the sandstones and to estimate their contribution to the gamma-ray spectrum. furthermore, the aim was to identify the gamma-spectral characteristics of intervals with different reservoir quality. with these objectives, we sampled the cores of west lulu-3 and defined three lithofacies based on petrographical and mineralogical data. the three lithofacies chosen are expected to have characteristic reservoir qualities as reflected by porosity and permeability. gamma-spectral data were obtained in the laboratory for each of the three lithofacies, and the laboratory data were compared with the log data. the harald field the harald field is situated in the danish portion of the søgne basin in the north-eastern part of the danish central graben (fig. 1). the hydrocarbon reservoir is located in sandstones of the middle jurassic bryne and lulu formations (johannessen & andsbjerg 1993; andsbjerg 2003, this volume; michelsen et al. 2003, this volume). according to johannessen & andsbjerg (1993) and andsbjerg (2003, this volume), this succession represents a range of paralic and coastal plain environments (fig. 2), and is overlain by the marine shales of the lola formation. the sng log over the cored section of west lulu-3 is represented in figure 3. geochemistry of potassium, thorium and uranium hassan et al. (1976) examined the mineralogy and chemical composition of 500 samples of varying lithology from different environments of deposition. of particular relevance to this work, their data included k, th and u measurements. the following conclusions concerning the occurrence of these three elements are based on hassan et al. (1976) as well as schlumberger (1982), and nielsen et al. (1987). potassium potassium is a major element in many rock-forming minerals. the radioactive isotope 40k constitutes 0.0118% of the total potassium in a mineral, and the gamma signal from the radioactive decay is an important indicator of mineralogical composition. the most common k-bearing minerals in sedimentary rocks are k-feldspar, mica and illite. k-feldspar and mica are more common in sandstones, illite in shales. k-feldspar and mica have a greater k content than illite, and thus the gamma-ray flux is only a measure of clay content in the absence of significant k-feldspar (cowan & myers 1988). glauconite can also be a significant source of k in sedimentary rocks. typical potassium contents of k-bearing minerals are: illite, 4.5%; muscovite, 7.9–9.8%; plagioclase, 0.54%; k-feldspar, 10.9–14.0% (nielsen et al. 1987). engstrøm (1981) measured k, th and u in danish cenozoic strata (clay, silt and sand) and found results in accordance with the distributions and concentrations mentioned above. kaolinite is sometimes reported to contain minor amounts of k and th, but ideal kaolinite contains neither elements (see below). straightforward relationships between k and clay content should not be expected. in miocene–pliocene shales from the north sea, berstad & dypvik (1982) found a positive correlation between k and the clay content. this was inferred to be mainly controlled by the illite of the clay fraction. on the contrary, in paleocene and eocene strata, radioactivity and clay content are negatively correlated. this is probably because the clay fraction is rich in a kpoor smectite, derived from basaltic volcanic material. thorium thorium is a common trace element in most geological environments. in weathering environments it is prac351 ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■ ■ ▲ ■■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■ ■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■■■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■ ■ w. lulu-4 w. lulu-3 feda graben inge highmid north sea high heno plateau tail end g raben arne–elin g raben ringkøbing–fyn high gertrud graben/ plateau outer rough basin coffee soil fault east n orth sea block søgne basin 56°30´n reverse fault triangles showing dip of fault plane normal fault boxes showing dip of fault plane salt structure well harald field fault plane ▲ ■■ 20 km 4°e 500 km uk norway denmark germany the netherlands fig. 1. the location of the west lulu-3 well within the harald field in the danish sector of the north sea central graben. inset map shows the central graben (blue) straddling the national boundaries and branching northwards into the outer moray firth and viking graben systems; the area of the enlarged map of the northern danish central graben is indicated in red. w. lulu-4 w. lulu-3 coal sequence boundary studied core section paralic and shallow marine (sand-dominated) fluvial and estuarine in incised valley alluvial plain and lacustrine shallow marine (mud-dominated) 50 m base jurassic 3784 m b. msl base jurassic 3793 m b. msl lola fm lulu fm bryne fm triassic gr sonicgr sonic ssw nne fig. 2. log correlation of the middle jurassic in the harald field area showing the stratigraphic position of the west lulu-3 cored section under study in this paper; for location of wells, see fig. 1. modified from andsbjerg (2003, this volume). 352 ii ii ii iii iii i i i iii iii iii iii iii iii iii ii ii i ii i i i iii shallow marine (lower shoreface) back barrier and shallow marine c oa st al p la in lo la f m lu lu f m br yn e fm estuarine and fluvial channels, peat mires alluvial plain li th os tr at ig ra ph y d ep th ( m b . m sl ) se di m en ta ry en vi ro nm en t li th of ac ie s sa m pl e a b c d1 d2 e f g h i j k1 k2 l1 l2 m n o p q1 q2 r s 3560 3580 3600 3620 3640 3660 3680 3700 3720 u (ppm) th (ppm) k (%) th/k (10-4) lithofacies from k-log -10 20 0 0 10 i ii iii10 0 5 tically insoluble and is thus commonly concentrated in residual deposits such as bauxite and clay. important quantities of thorium are found in the heavy minerals monazite, rutile and zircon. zircon normally contains 100–2500 ppm th (nielsen et al. 1987), while monazite may contain substantially more. pure clay samples typically contain 5–30 ppm th (adams & weaver 1958). theoretical considerations indicate that th4+, with an ionic radius of 0.97 nm, cannot be accommodated in the layer structure of clay minerals, so that th in these minerals could be assumed to be fixed by adsorption (adams & weaver 1958). hurst & milodowski (1994) stated, however, that the origin of th in clay minerals is unclear, and proposed that it is caused by inclusions of clay-sized heavy minerals. in the samples of hassan et al. (1976), the th content correlated with the content of clay minerals, but it was not proven that the clays were responsible for the th enrichment. uranium under neutral ph conditions, the uranyl ion ((uo)2+) forms ionic complexes with carbonates, which control its dispersion and mobility in nature. the uranyl ion also forms numerous complexes with organic compounds (e.g. humic acids), which facilitate its fixation by organic and mineral matter. hassan et al. (1976) found that u shows a strong correlation with organic carbon, probably because uo2 precipitates under reducing conditions. factors other than the content of organic matter (such as the availability of u) may be important for the occurrence of u in sediments as exemplified by the data of berstad & dypvik (1982). in shales from the cenozoic of the north sea, these workers found no correlation between u and total organic carbon (toc). uranium also occurs as a trace element in accessory minerals (zircon contains 300–3000 ppm u; nielsen et al. 1987) and in colloidal fe-oxide/hydroxide coatings on mineral grains. uranium species may also be adsorbed onto clay minerals. methods borehole gamma-ray measurement the sng tool used for logging of the west lulu-3 borehole uses five continuous energy windows covering the entire energy range from 0.1–3.0 mev (fig. 4). the heart of the tool is a sodium iodide (nai) crystal that detects the gamma rays in the borehole. a gamma-ray detector based on a nai crystal is able to produce accurate data on the radioactivity in the surroundings. to interpret the spectrum correctly, however, it is necessary to take into account the influence of the drilling mud, variations in the borehole diameter and the standoff of the probe. the (net) uranium count rate in window w4 is usually low (fig. 4), and the uranium measurements therefore become dependent on the count rates of the low-energy windows w1 and w2, which are influenced the most by the borehole parameters. this results in an unreliable u log curve. by using the whole spectrum, high count rates are attained, i.e. the statistical accuracy of the counting is high, but borehole correction factors must be included. laboratory gamma-ray measurement similar techniques can be used for laboratory measurements of gamma radiation from k, th and u in rock samples. laboratory measurements were obtained with nai crystals using three energy windows centred around the three most prominent gamma energy k, th and u peaks of the spectrum (fig. 4). these measurements are in the high-energy part of the spectrum and thus less dependent on geometry, density and matrix influence. however, by using only counts from a fraction of the spectrum, longer counting times are needed. for the present study, 19 core-sections (each c. 80 cm long) of the west lulu-3 cores were selected on the basis of core inspection and the k, th and u signals of 353 facing page: fig. 3. the natural spectral gamma-ray log over the cored interval of west lulu-3. samples investigated in the present study are indicated by the letters a–s. for the construction of the lithofacies log, the k value from the spectral gamma borehole log was multiplied by 0.8 (as recommended in the text). intervals in which the k content is < 0.7% are referred to lithofacies i, those with 0.7–2% k are referred to lithofacies ii and the sediments with > 2% k are assigned to lithofacies iii. the corresponding lithofacies of the core samples, as determined by laboratory measurements, are indicated by roman numerals. the sedimentary environment interpretations are by j. andsbjerg (personal communication 1995; see also andsbjerg 2003, this volume). 354 the sng log. ten of the sections are from the bryne formation, six are from the lulu formation and the upper three sections are from the lower levels of the lola formation (figs 2, 3). the sections were chosen to cover the range of different k, th and u concentrations and ratios in the interval 11 800–12 300 ft below kelly bushing , corresponding to c. 3560–3720 m below mean sea level. each section was measured in the laboratory by a spectral natural gamma-ray scanner with a nai crystal, and on the basis of these results, typical core intervals, 5–7 cm long, were selected from each core section for more detailed analysis. a total of 23 samples were collected. for each of these samples, the k, th and u contents were determined by a ge(li) gamma spectrometer. most samples were selected from intervals with stable k, th and u readings. mineralogy and chemistry after removal of a reference slab, samples were taken for the preparation of thin sections and polished specimens. the thin sections were described by optical microscopy, and the polished specimens were used for back-scatter electron microscopy (bsem) and energydispersive microprobe analysis (eds). for each selected interval, a sample of c. 125 g was cut out, crushed and homogenised. measurement of toc, inorganic carbon and sulphur was undertaken on 0.5 g size sub-samples after ignition in a leco cs 225 furnace. the k, th and u content of the crushed samples was measured by placing the samples inside a laboratory nai crystal gamma spectrometer. by this method, low levels of radioactivity can be measured. approximately 120 g of each sample were subsequently separated into grain-size fractions by wet and dry sieving. the mineral content of the clay size (< 2 µm), fine silt size (2–20 µm), coarse silt – very fine sand (20–125 µm), and coarser sand fractions (125–250 µm, 250–600 µm and > 600 µm) was determined by x-ray diffractometry (xrd) using ni-filtered cu-kα radiation at apparatus settings of 40 kv and 40 ma with the step scanning variant slit. the effectiveness of the grain-size separation and the mineralogy of the fractions was checked by microscopy. on the basis of these data, the separated fractions for each sample were remixed into three portions: clay (< 2 µm), silt and very fine sand (2–125 µm) and coarser sand (> 125 µm). these three portions were chosen because within each portion, only minor differences in mineralogy were observed. for each of the three portions, the k, th and u concentrations were determined by gamma-ray spectrometry. separation of the samples into fractions of individual minerals was not attempted, but the mineralogical composition of the samples was estimated, and the contribution from each mineral to the total k was inferred in the following way. the amount of organic matter in each sample was assumed to be equal to the toc. the amount of carbonate was calculated from inorganic carbon with reference to xrd data. all sulphur was assumed to be associated with pyrite unless barite was detected, in which case optical microscope and xrd data formed the basis for partitioning the sulphur between barite and pyrite. the amounts of k-feldspar, w1 w2 w3 w4 w5 1 mev0 10 0 00 c ou nt s pe r ch an ne l ( 10 00 s ec ) 20 0 00 2 mev energy k window u window th window fig. 4. natural gamma-ray spectrum measured in the laboratory with a nai crystal. energy windows used for concentration determinations are centred around the three most prominent highenergy k, th and u peaks. natural gamma-ray spectra measured in boreholes contain the same peaks, but the spectra are blurred compared to a laboratory-measured spectrum, and the peaks are less distinct. the tool for the natural gamma ray spectrometry log in the west lulu-3 well uses five continuous windows, w1–w5, covering the energy range from 0.1 to 3.0 mev. at lower energies, the spectrum shape is significantly influenced by the borehole geometry and mud density. 355 clay minerals and mica were calculated from the gamma spectrometry and eds potassium measurements, guided by the semi-quantitative xrd and microscope data on mineral abundance. on the basis of these data and eds analysis of the k content in the observed minerals, the contribution from each mineral to the total k was calculated. the presence and identity of heavy minerals were determined by microscopy and eds analysis. results facies analysis on the basis of thin-section microscopy, the samples were classified into lithofacies i, ii and iii; a similar classification was developed by vernik & nur (1992; table 1). lithofacies i (five samples) comprises well-sorted quartz arenites with a maximum clay content of 2–3%. the samples are from sedimentary rocks that are primarily interpreted to represent channel-fill sediments within paralic and coastal plain/floodplain environments (fig. 3; johannessen & andsbjerg 1993; andsbjerg 2003, this volume). porosity is in the range 23–28%, and permeability is in the range 200–2000 md. lithofacies ii (eight samples) consists of grain-supported quartz arenites with up to 18% clay. the samples are from sedimentary rocks interpreted to represent coastal plain/floodplain and lower shoreface environments of deposition (fig. 3; johannessen & andsbjerg 1993; andsbjerg 2003, this volume). porosity is in the range 11–17% and permeability is in the range 0.4–25 md. lithofacies iii (ten samples) comprises wackes and sandy shales (matrix-supported, with occasional grainto-grain contacts) containing more than 18% clay. the 93 1 1 1 1 3 87 1 1 5 1 1 3 1 89 1 1 5 1 3 93 3 1 1 1 1 81 1 1 13 3 1 88 2 1 7 1 1 88 2 1 7 1 1 88 2 1 3 3 3 73 3 2 2 5 11 1 1 2 77 2 3 9 9 68 6 24 2 59 2 1 15 15 8 60 3 4 9 18 2 2 2 60 3 4 9 18 1 2 2 1 38 1 16 33 2 10 41 1 10 45 3 43 4 15 30 8 32 1 6 56 5 44 1 2 25 25 3 25 1 1 31 16 2 23 1 28 14 56 1 1 36 1 7 35 18 2 1 27 1 2 45 23 1 1 h 3619.4 i 0.09 0.26 0.4 e 3588.5 i 0.14 2.47 0.1 g 3610.4 i 0.03 0.64 0.1 m 3665.2 i 0.13 0.13 0.3 j 3635.6 ii 0.04 3.18 0.0 l1 3660.7 i 0.08 0.25 0.2 l2 3661.0 i 0.13 0.19 0.1 s 3720.1 i 0.38 0.08 0.0 a 3561.2 ii 1.40 0.67 0.5 q2 3698.7 ii 0.03 0.19 0.0 k2 3654.2 iii 0.04 1.88 0.1 r 3700.0 ii 0.05 8.23 0.2 b 3566.3 ii 0.01 2.18 1.0 c 3573.9 ii 0.10 2.20 0.9 d2 3585.9 iii 0.31 10.25 0.1 f 3591.6 iii 0.03 2.84 0.1 d1 3585.5 iii 0.05 8.28 0.2 i 3627.0 iii 0.00 5.00 0.1 q1 3698.0 iii 0.32 0.12 0.0 p 3689.7 iii 2.70 0.93 0.0 k1 3653.8 iii 0.10 1.10 0.1 n 3671.0 iii 0.30 0.80 0.0 o 3680.7 iii 0.02 0.51 0.0 0.03 0.0 0.1 0.15 0.15 0.07 0.0 0.15 0.15 0.07 0.0 0.15 0.03 0.0 0.07 0.4 0.4 0.07 0.0 0.15 0.2 0.0 0.1 0.3 0.2 0.0 0.1 0.3 0.03 0.15 0.1 0.3 0.01 0.14 0.2 0.44 0.07 0.09 0.5 0.3 0.3 0.7 0.4 0.0 0.0 0.15 0.8 0.1 0.3 0.04 0.7 0.4 0.44 0.07 0.04 0.7 0.4 0.44 0.018 1.0 1.1 0.1 0.0 1.4 0.7 0.0 0.15 0.9 1.1 0.4 0.0 1.7 0.4 0.1 0.0 0.3 1.3 0.2 0.15 0.16 2.2 0.1 0.15 1.7 1.0 0.0 0.0 0.18 2.5 0.7 0.15 0.3 3.2 0.2 0.15 in or ga ni c ca rb on (c ar bo na te ) o rg an ic c ar bo n (t o c ) su lp hu r qu ar tz ( + a lb ite ) k -fe ld sp ar m us co vi te ill ite ka ol in ite do lo m ite si de ri te py ri te co al ba ri te he av y m in er al s gl au co ni te ka ol in ite ill ite m us co vi te fe ld sp ar gl au co ni tesample* geochemical data (wt%) interpreted mineral content (%; based on xrd, bsem and petrography) interpreted contribution to k content (%) depth metre b. msl lithofacies * samples listed in order of increasing k content (bulk sample), see table 3. table 1. geochemical data, mineralogy and interpreted mineral contribution to k content sampled facies are interpreted, in the main, to represent a coastal plain/floodplain environment of deposition (fig. 3; johannessen & andsbjerg 1993; andsbjerg 2003, this volume). porosity is less than or equal to 11%, and permeability below 0.5 md. characteristic bsem images of each lithofacies are presented on figure 5. crushing and grain size separation were successful for samples of lithofacies i, and some of lithofacies ii, but not for samples of lithofacies iii and several lithofacies ii samples, where the coarser fractions contain mineral aggregates. although the fine fractions may include pieces broken from the coarser grains during sample crushing, no such pieces were observed. the sieve data are presented in table 2. mineralogy quartz is present in all samples as sandand silt-sized grains. samples of lithofacies i contain sand-size quartz grains with diameters greater than 130 µm, samples of lithofacies ii have quartz grains of very fine sand and coarse silt size (50–130 µm), and samples of lithofacies iii, as a rule, contain only silt-size particles. one sample of lithofacies iii though, contains quartz grains as 356 0,1 mm 0,1 mm 0,1 mm 1 2 3 fig. 5. back-scatter electron microscope images. a: lithofacies i, core sample e. porous, grain-supported quartz arenite with isolated heavy mineral grains, such as zircon (1). b: lithofacies ii, core sample c. grain-supported quartz wacke, with no visible porosity. grains include: 2, pyrite; 3, glauconite. c: lithofacies iii, core sample k2. matrix-supported quartz wacke. sample‡ lithofacies (> 125µm) (2–125µm) (< 2µm) total (%) (%) (%) (%) h i 77.99 17.63 3.18 98.8 e i 49.08 46.41 3.04 98.53 g i 69.02 24.44 2.84 96.3 m i 58.2 37.64 2.71 98.55 j ii 45.19 46.52 6.76 98.47 l1 i 64.07 31.08 3.94 99.09 l2 i 62.11 32.39 6.47 100.97 s i 74.05 21.57 3.85 99.47 a ii 0.6 89.67 6.78 97.05 q2* ii 25.97 13.95 k2* iii 9.8 70.37 18.54 98.71 r* ii 7.51 81.96 7.3 96.77 b* ii 3.07 83.48 8.61 95.16 c* ii 4.24 82.82 10.49 97.54 d2* iii 24.22 63.8 7.45 95.47 f* iii 23.79 58.95 15.72 98.46 d1* iii 23.69 66.38 6.8 96.87 i* iii 21.68 62.86 11.32 95.86 q1* iii 10 73.92 17.4 101.32 p* iii 31.28 54.7 10.05 96.03 k1* iii 28.89 52.31 13.71 94.91 n* iii 14.12 67.84 9.6 91.56 o* iii 16.62 65.62 17.26 99.5 *the silt and sand fractions contain lumps of non-disaggregated material. ‡ samples listed in order of increasing k content (bulk sample), see table 3. table 2. grain-size fractions of crushed samplesa b c large as 80 µm. some diagenetic quartz overgrowths are seen in samples of lithofacies i and ii. feldspar. k-feldspar grains have been detected in 18 of the 23 samples from a combination of xrd, thin-section study and eds. in 16 of the samples, grains of plagioclase and k-feldspar have been found, and in four samples only plagioclase was found. the only type of plagioclase detected by eds is albite. the content of plagioclase barely exceeds a few percent by volume in any sample, and no quantification was attempted. the estimated quartz content (table 1) therefore probably includes a minor component of plagioclase. both degraded and fresh feldspars are observed in lithofacies iii samples, whereas feldspar grains in lithofacies i and ii samples are generally degraded. no indications of preferential degradation of either plagioclase or k-feldspar was noted. muscovite was found in most samples (table 1). the textural relationships indicate that the muscovite is of detrital origin. it is generally fresh and unaltered. eds analysis indicated a k content close to the ideal (10%). illite is present in the clay fraction of most samples, with the exception of some from lithofacies i (table 1). eds analysis indicated that the illites of samples a–p contain around 7% k and the illites of samples q–s contain 5% k. illite was found in clay of apparently detrital origin (fig. 6a). 357 0,1 mm 0,1 mm 0,1 mm 0,1 mm fig. 6. back-scatter electron microscope images. a: allochthonous clay (illite and kaolinite) between sand grains in sandstone of lithofacies ii; core sample j. b: authigenic kaolinite platelets in sandstone of lithofacies i. the bright grain in the upper part of the picture is rutile; core sample m. c: particles of barite (white) caught in the pore space of a lithofacies i sandstone, due to drilling mud invasion; core sample h. d: the fabric observed in c, at higher magnification, showing the characteristic textural habit of barite. a c d b 358 kaolinite was detected in all samples, being the dominant mineral in several samples of lithofacies iii. authigenic kaolinite as well as detrital kaolinite were observed (fig. 6). no attempt was made to differentiate between different polytypes. the measured k content of the kaolinites (by eds) in the present samples have the following general pattern: samples a–c, around 0.1% k; samples d–m, approximately 3% k; samples n–s, approximately 1% k. heavy minerals. grains of zircon, rutile and chromite were identified by optical microscopy and eds. they occur in significant amounts in sample e and g of lithofacies i, and in sample j of lithofacies ii. pyrite is present in most of the samples where it is commonly associated with organic matter (fig. 5). carbonate. dolomitic cement was found in several samples (table 1). in sample a, fe-rich dolomite was identified with a mg/fe (atomic) ratio of 1.4 (based on eds). in sample p (lithofacies iii), many siderite concretions, as well as dolomite, are present. the illite-rich sample p is from an interval that is interpreted to record an alluvial plain setting (fig. 3; andsbjerg 2003, this volume). this concurs with the common observation of siderite in anoxic freshwater deposits (postma 1983). barite is present in minor amounts, as indicated by the xrd pattern of four samples of lithofacies i (table 1). barite was only found in the clay and silt fractions. bsem imaging reveals the textural habit of the barite, showing that the barite content is the result of infiltration of fine-grained particles from the drilling mud into permeable lithofacies i sandstones (fig. 6c, d). glauconite was found in the three samples (a–c) from the lola formation. this is thought to concur with the interpreted environment of deposition (lower shoreface). a potassium content of 2–4 wt% was measured by eds. gamma-spectral analysis the results of the gamma-spectral analysis of the crushed samples are listed in table 3. the results from gammaspectral analysis of the clay fractions are given in table 4. gamma-spectral analysis of clay fraction sample* lithofacies k (%) th (ppm) u (ppm) h i 1.65 ± 0.10 7 ± 3 2.4 ± 0.6 e i 1.80 ± 0.14 18 ± 4 1.7 ± 0.8 g i 2.28 ± 0.13 9 ± 3 1.9 ± 0.8 m i 2.81 ± 0.15 15 ± 4 3.5 ± 0.9 j ii 2.36 ± 0.07 8.7 ± 1.5 2.7 ± 0.3 l1 i 2.50 ± 0.10 15 ± 3 2.1 ± 0.6 l2 i 2.02 ± 0.07 10.9 ± 1.7 1.5 ± 0.4 s i 2.98 ± 0.12 12 ± 3 1.6 ± 0.7 a ii 3.50 ± 0.11 18 ± 2 1.6 ± 0.5 q2 ii 2.88 ± 0.05 10.1 ± 1.1 2.1 ± 0.3 k2 iii 1.63 ± 0.05 12.3 ± 1.1 3.6 ± 0.2 r ii 4.05 ± 0.07 19.0 ± 1.4 5.9 ± 0.3 b ii 2.98 ± 0.08 13.8 ± 1.7 3.5 ± 0.4 c ii 3.54 ± 0.05 15.2 ± 1.0 4.1 ± 0.2 d2 iii 2.14 ± 0.06 10.9 ± 1.4 2.4 ± 0.3 f iii 2.41 ± 0.05 14.2 ± 1.0 3.8 ± 0.2 d1 iii 2.64 ± 0.06 12.0 ± 1.4 4.7 ± 0.3 i iii 2.42 ± 0.05 13.1 ± 1.0 3.0 ± 0.2 q1 iii 3.41 ± 0.05 12.0 ± 0.1 2.1 ± 0.2 p iii 4.17 ± 0.08 18.5 ± 1.6 3.1 ± 0.4 k1 iii 3.36 ± 0.07 15.7 ± 1.3 3.9 ± 0.3 n iii 4.35 ± 0.06 16.5 ± 1.2 2.4 ± 0.3 o iii 4.29 ± 0.05 16.2 ± 0.9 2.7 ± 0.2 uncertainties are one standard deviation. *samples listed in order of increasing k content (bulk sample), see table 3. table 3. gamma-spectral analysis of bulk sample sample* lithofacies k (%) th (ppm) u (ppm) h i 0.25 ± 0.01 1.0 ± 0.4 0.4 ± 0.1 e i 0.42 ± 0.02 8.2 ± 0.5 3.2 ± 0.1 g i 0.47 ± 0.02 3.8 ± 0.4 1.5 ± 0.1 m i 0.57 ± 0.01 2.7 ± 0.4 1.1 ± 0.1 j ii 0.59 ± 0.01 1.8 ± 0.4 0.3 ± 0.1 l1 i 0.60 ± 0.02 2.6 ± 0.6 0.5 ± 0.1 l2 i 0.60 ± 0.02 5.5 ± 0.4 0.9 ± 0.1 s i 0.67 ± 0.02 1.0 ± 0.4 0.1 ± 0.1 a ii 0.87 ± 0.02 3.9 ± 0.5 1.2 ± 0.1 q2 ii 1.18 ± 0.02 5.7 ± 0.5 1.2 ± 0.1 k2 iii 1.21 ± 0.03 10.7 ± 0.6 3.9 ± 0.1 r ii 1.92 ± 0.03 8.0 ± 0.6 2.6 ± 0.1 b ii 1.99 ± 0.03 8.0 ± 0.6 2.8 ± 0.1 c ii 2.00 ± 0.03 11.6 ± 0.6 3.2 ± 0.1 d2 iii 2.05 ± 0.04 13.3 ± 0.7 4.0 ± 0.2 f iii 2.15 ± 0.03 13.2 ± 0.6 4.0 ± 0.1 d1 iii 2.17 ± 0.02 10.2 ± 0.5 3.4 ± 0.1 i iii 2.17 ± 0.03 11.1 ± 0.7 3.3 ± 0.1 q1 iii 2.37 ± 0.03 10.3 ± 0.6 2.5 ± 0.1 p iii 2.42 ± 0.03 7.3 ± 0.5 2.7 ± 0.1 k1 iii 2.89 ± 0.03 12.3 ± 0.6 2.7 ± 0.1 n iii 3.30 ± 0.03 16.2 ± 0.6 3.6 ± 0.1 o iii 3.51 ± 0.03 16.9 ± 0.6 3.2 ± 0.1 sample size: 18 cm3. uncertainties are one standard deviation. *samples listed in order of increasing k content (bulk sample). table 4. the data for the coarser fractions are not listed. these fractions do not represent the values for the ‘true’ grain size interval because of the incomplete disaggregation of the original sample. the interpreted total mineralogy of the samples, and the interpreted contribution from each mineral to the total k content of each sample, are listed in table 1. potassium concentrations are low in lithofacies i samples (< 0.7%), higher in lithofacies ii samples (0.6–2%), and highest in lithofacies iii (generally > 2% although 359 4 3 2 1 0 20 15 10 5 0 5 4 3 2 1 0 25 20 15 10 5 0 5 4 3 2 1 0 8 6 4 2 0 3550 3575 3600 3625 3650 depth (m b. msl) depth (m b. msl) depth (m b. msl) depth (m b. msl) depth (m b. msl) depth (m b. msl) 3675 3700 3725 3550 3575 3600 3625 3650 3675 3700 3725 3550 3575 3600 3625 3650 3675 3700 3725 3550 3575 3600 3625 3650 3675 3700 3725 3550 3575 3600 3625 3650 3675 3700 3725 3550 3575 3600 3625 3650 3675 3700 3725 k in t ot al s am pl e (% ) t h in t ot al s am pl e (p pm ) u in t ot al s am pl e (p pm ) k in c la y fr ac tio n (% ) t h in c la y fr ac tio n (p pm ) u in c la y fr ac tio n (p pm ) lithofacies i lithofacies ii lithofacies iii fig. 7. content of k, th and u in total crushed samples and in clay fractions versus sample depth. samples of lithofacies i contain less than 0.7% k, whereas the th and u values are variable, dependent on a varying content of heavy minerals. samples of lithofacies ii have 0.6–2% k, and variable contents of th and u. samples of lithofacies iii generally have more than 2% k, more than 7 ppm th and more than 2.5 ppm u. the clay fraction data indicate a stratigraphic variation in k content, decreasing slightly with depth down to 3650 m followed by a broad increase between 3650 and 3725 m. the th content tends to follow the same pattern as k. similar variations are not apparent for u. 360 one sample contains 1.2% k; figs 7, 8). this differentiation of lithofacies in terms of k content was not observed for the clay fraction. considering all the data, the k content of the clay component shows a weak decrease with increasing depth to around 3650 m; at greater depths, the data show considerable scatter but appear to broadly increase between 3650 and 3725 m (fig. 7). the th and u concentrations of the total samples are variable in lithofacies i and ii, but uniformly high in lithofacies iii (fig. 7). the th content in the clay fraction follows roughly the same pattern as the k content (see above), while no pattern is apparent for the u content of the clay fraction (fig. 7). for lithofacies ii and iii samples, roughly constant th/k ratios of 3 x 10-4 – 10 x 10-4 were obtained, whereas lithofacies i samples show significant variation in this ratio (2 x 10-4 – 22 x 10-4; fig. 9). this is a consequence of lithofacies i samples being low in k but having variable th and u contents; it is assumed that this variability is governed by the heterogenereous distribution of heavy minerals within the facies. total organic carbon (toc) fragments of organic matter are found in most samples of lithofacies ii and iii and in a single sample of lithofacies i. a general negative relationship between car250 200 150 100 50 0 0 1 2 k (%) 3 4 q ua rt z gr ai n si ze ( µm ) lithofacies i lithofacies ii lithofacies iii fig. 8. median size of quartz grains, estimated from thin sections, vs. k content. samples of lithofacies i contain less than 0.7% k and comprises quartz grains in the range 130–220 µm. samples of lithofacies ii contain 0.6–2% k and consist of quartz grains in the range 50–130 µm. samples of lithofacies iii generally contain more than 2% k and are made up of quartz grains less than 80 µm in size. fig. 9. th/k ratio vs. k content. the th/k ratio does not effectively separate the lithofacies i, ii and iii, although lithofacies i samples rich in heavy minerals can be differentiated by their high th/k ratio. 0.0025 0.0020 0.0015 0.0010 0.0005 0 0 1 2 3 4 t h/ k r at io k (%) lithofacies i lithofacies ii lithofacies iii fig. 10. u concentrations in the total samples and in the clay fractions vs. toc. no general relationship between u and toc is apparent from the two figures. total samples of lithofacies i are low in toc and show a large variation in u, samples of lithofacies ii show a scattered pattern, while samples of lithofacies iii give constant high u values and a large variation in toc. 5 4 3 2 1 0 0 2 4 6 8 10 12 u in t ot al s am pl e (p pm ) 8 6 4 2 0 0 2 4 6 8 10 12 u in c la y fr ac tio n (p pm ) toc (wt%) toc (wt%) lithofacies i lithofacies ii lithofacies iii bonate and toc can be noticed (table 1). samples of lithofacies i give low toc values but show a large variation in u content (fig. 10). samples of lithofacies ii display a large variation in both toc and u, while lithofacies iii samples have a high content of u and a large toc variation. uranium concentration and toc are apparently independent. location of potassium the data suggest that feldspar is the main source of k in the k-poor lithofacies i sandstones, while mica and kaolinite are a minor source (fig. 11). the intermediate k content of lithofacies ii sandstones is probably primarily due to feldspar, mica and illite, while kaolinite contributes less. in the k-rich lithofacies iii samples, feldspar is of only minor importance, whereas illite and kaolinite are inferred to have roughly equal importance as a source of k. in figure 11, inferred contributions to the total k from the minerals in each sample are presented in order of depth. the low k values measured by eds in kaolinites from below 3660 m (i.e. samples l1–s) are responsible for the low contribution to the total k attributed to kaolinite below this depth. 361 a b c d1 d2 e f g h i j k1 sample k/kmax k2 l1 l2 m n o p q1 q2 r s 0.0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1.0 i ii iii a b 4 3 2 1 0 k in m in er al s (% ) 4 3 2 1 0 k in m in er al s (% ) glauconite feldspar muscovite illite kaolinite fig. 11. the cumulative contributions of the k-bearing minerals to the total k concentration, both for individual samples (a) and ranked in progressive k content (b). glauconite is only present in the samples from the lola formation (above 3575 m, i.e. samples a–c), while illite gains drastically in importance relative to kaolinite below 3665 m (samples n–s). in samples of lithofacies i, k-feldspar is the dominant source of potassium; in samples of lithofacies ii, k-feldspar, muscovite and illite dominate; in samples of lithofacies iii, illite and partly illitised kaolinite are the major sources of potassium. 362 laboratory gamma spectrometry of cores vs. log data comparison between the sng log data and the results from the measurements of the gamma radiation in the laboratory reveals significant differences (fig. 12). on average, the laboratory data indicate 20% lower k concentrations than the log data. for th, the laboratory results are on average 5% lower than the log values. interpretation of the u data is complicated; the log shows negative concentrations in some intervals, and no correlation was found between log data and laboratory measurements. this concurs with the findings of hurst (1990) who also observed that u values from the sng log in north sea boreholes were unreliable. discussion diagenetic factors the coexistence of kaolinite and feldspar in sandstones has been the subject of much discussion in the literature (lindgreen 1985; bjørkum & gjelsvik 1988; aja et al. 1991; bjørkum et al. 1993). at temperatures higher than 50–80°c (or according to bjørlykke & egeberg (1993), above 90–100°c), kaolinite and k-feldspar may react to form authigenic illite and quartz cement. the extensive illitisation of sandstones of the garn formation, offshore norway, is interpreted to have arisen because the sandstone reached either a temperature of 140°c or a time-temperature index (tti) of 100 (ehrenberg 1990; see waples 1980 for an introduction to the tti maturity index). a temperature near the critical value of 140°c was recorded during logging of the studied reservoir interval of west lulu-3. quartz cement is frequently observed, and some detrital kaolinite contains up to 3% k, while no k was detected in the authigenic kaolinite. an explanation for the varying degree of illitisation of kaolinite can be a difference in polytype between the detrital and authigenic kaolinite. ehrenberg et al. (1993) described replacement of kaolinite by dickite in three sandstone formations on the norwegian continental shelf. apparently, the transformation from kaolinite to dickite takes place at temperatures above 120–130°c via dissolution/re-precipitation. ehrenberg et al. (1993) found more extensive illitisation of kaolinite than of dickite, possibly because this process is kinetically favoured over illitisation of dickite. the presence of carbonate and kaolinite cements is thought to reflect the release of acid pore waters from shales. milliken & land (1991) related the acid pore waters to illitisation of smectite, whereas curtis (1983) and hansley & nuccio (1992) attributed the acid pore water to the maturation of kerogen. the acid pore water destabilises carbonate and aluminosilicates in the shales, co2 and metal ions go into solution, and sediment compaction forces some of these modified pore waters into sandstones. if the sandstones contain reactive feldspar (i.e. detrital metamorphic or magmatic feldspars), acids 3654 3656 3658 3660 depth (m b. msl) 3662 3664 3666 3668 4 2 0 20 10 0 4 2 0 k ( % ) li th ofa ci es t h (p pm ) u ( pp m ) k1k2 l1l2 m samples iii ii i ii lab. data log data fig. 12. comparison between the log data and the results from the measurements in the laboratory for core sections k, l and m. the k and th data compare reasonably well, whereas the u readings are unrelated. the lithofacies derived from the log values are marked by roman numerals. from the shales are effectively buffered, thus allowing precipitation of carbonate and/or kaolinite. curtis (1983) concluded that once the supply of reactive detrital feldspar is exhausted through dissolution and albitisation, shale-derived hydrogen ions remobilise carbonate in sandstones, generating secondary porosity and fluids with high pco2. no secondary porosity was recognized in thin sections of the west lulu-3 samples, and in line with the interpretation of curtis (1983), feldspar is present in practically all samples. lithofacies identification based on gamma spectrometry as demonstrated by figures 7 and 8, the samples from the three lithofacies are most effectively separated on the basis of the k content. thus grain-supported porous (clay-poor) quartz arenites (lithofacies i) contain < 0.7% k, grain-supported, clay-rich quartz arenites (lithofacies ii) contain 0.6–2% k whereas the matrix-supported wackes of lithofacies iii contain > 2% k. in figure 3, the distribution of these three lithofacies is presented, calculated on the basis of (corrected) log data. lithofacies ii occurs throughout the section but particularly dominates the uppermost interval where it is interpreted to represent a lower shoreface sedimentary environment (j. andsbjerg, personal communication 1995). lithofacies i, showing good reservoir properties, is common in the upper part of the bryne formation and in the lulu formation where it typically occurs within channel-fill sections of paralic (estuarine) origin (j. andsbjerg, personal communication 1995; andsbjerg 2003, this volume). the clay-dominated lithofacies iii is mainly found in the lower part of the studied section which is interpreted to represent an alluvial plain sedimentary environment (j. andsbjerg, personal communication 1995; andsbjerg 2003, this volume). schlumberger (1982) reported how the nsg tool can differentiate between different types of sandstone. feldspathic sandstones or arkoses will typically have th/k ratios less than 1 x 10-4, because of their low th and their high k contents whereas micaceous sandstones have a th/k ratio close to 2.5 x 10-4 because of the presence of th-bearing heavy minerals. the th/k ratios in claybearing sandstones are still higher. this is consistent with the present clay-bearing samples (lithofacies ii and iii) which have th/k ratios of 3 x 10-4 – 10 x 10-4. by contrast, this ratio varies widely in the lithofacies i samples (th/k: 2 x 10-4 – 22 x 10-4), depending on the feldspar and heavy mineral content. thus, high th/k ratios can be found in mineralogically mature lithofacies i sandstones. the lithofacies as defined in this study are expected to reflect the reservoir quality of the rock, but they cannot be predicted from the th/k ratio. hurst (1990) also found th/k cross-plots to be poorly suited to identifying and quantifying the clay mineralogy of sandstones. potassium content of minerals as noted above, feldspar is inferred to be the main source of k in lithofacies i and ii samples, whereas illite and kaolinite are the dominant sources of k in lithofacies iii samples. the k-feldspars of lithofacies i and ii samples commonly appear degraded, but a k content of c. 14% is typical. the illite has a highly variable k content, whereas the larger mica grains (muscovite) have a k content close to the ideal (10%). some kaolinite grains (mainly diagenetic) are free of k, while others (probably detrital) have a k content of about 3%. in the samples from the shallow marine lola formation, glauconite contributes to the k content. the k content of the glauconite was measured to be 2–4% by eds. the k content of the clay fraction of the samples shows no apparent relationship with lithofacies, but shows broad variation with depth, as described earlier. this could be explained by a change in clay mineralogy due to diagenesis (see above), as some indications of transformation of kaolinite into illite were noted. this explanation, however, is contradicted by two observations: (1) the pattern does not show a uniform depthrelated trend and (2) several of the deepest samples are rich in kaolinite relative to illite. the best explanation for the broad variations with depth is probably stratigraphic variations in the provenance of the clastic material. sources of uranium and thorium hassan et al. (1976) and schlumberger (1982) have reported that uranium commonly shows a strong correlation with organic carbon. in common with the data of berstad & dypvik (1982), this study has demonstrated no such correlation. imam & trewin (1991) stated that in the sandstones of the claymore sandstone member of the kimmeridge clay formation, the contribution to radioactivity from organic matter is minor, while the high gamma-ray radiation is considered to be due to high k-feldspar concentrations and radioactive heavy minerals, particularly zircon and monazite. our data support their conclusions. 363 comparison of log and laboratory data the observed discrepancy between log and laboratory measurements could be due to calibration problems, differences in measuring geometry or due to the influence of the drilling mud. calibration one may assume that both borehole tool and laboratory equipment are in this case calibrated correctly. the laboratory equipment for measurements of samples is regularly checked using reference material from the international atomic energy agency in vienna (iaea 1987). the core scanner is checked using secondary reference material. according to standard practice, the sng tool is calibrated regularly. measuring geometry the individual samples investigated in the laboratory are much smaller than the volume contributing to the measurements of the borehole logging tool in the borehole. however, systematic variations such as those observed for k and th cannot be explained in this way. moreover, most samples were selected from intervals with stable concentrations of th, u and k, so that this source of error is considered to be of minor importance. the crushed samples were measured in a fixed geometry (in a laboratory nai crystal detector), and the core slices of different thickness were measured using reference material with similar geometries. the accuracy of the core scanner is dependent on core geometry (regular or fractured cores). for the cores investigated, such influence is minor and, in any event, for irregular cores the concentrations of k, th and u should be influenced to almost the same extent. the accuracy of borehole measurements is dependent on borehole geometry (variations in diameter). however, the associated caliper log indicates a fairly constant borehole diameter. borehole mud borehole measurements are influenced by the density and chemistry of the drilling mud. no major potassium-, thorium-, or uranium-bearing additives (e.g. kcl) were present in the mud. major amounts of barite were added in order to achieve a mud density of 1.9 g/cm3. barite attenuates gamma-rays entering the borehole from the formation, especially in the lower part of the energy spectrum. if this influence is not properly taken into account, low or even negative u concentrations may result from the calculations. from data reported by ellis (1982), calculated th concentrations may be expected to be higher than the real concentrations under conditions of high barite content. normally, a correction for high density borehole mud is included in the calculations by the logging company, but that apparently was not the case for the west lulu-3 log. the problem is not trivial because it is impossible to perform recalculations without detailed information on the correction program. by multiplying the log concentrations of th and k by 0.95 and 0.80 respectively, values approximately similar to those measured in the laboratory can be obtained. unfortunately the log concentrations of u cannot be transformed to comply with the laboratory values by any simple procedure. conclusions 1. from optical microscopy, x-ray diffraction and eds microprobe data, we infer that the main sources of potassium are feldspar, muscovite, illite and illitised kaolinite. k-feldspar is the main source in samples with a limited amount of clay, while illite and illitised kaolinite are most important in the clay-rich intervals. 2. neither gamma-spectral k nor th was found to be a measure of clay content, but the k concentration provides a method of recognising the three broad lithofacies adopted here. porous quartz arenites have k concentrations below 0.7%, grain-supported quartz arenities with pore-filling clay have k concentrations between 0.6 and 2%, while matrix-supported wackes normally have k concentrations above 2%. it is possible, therefore, to define the lithofacies on the basis of the k signal of the spectral natural gamma log. 3. the k content of the clay fraction varies as a function of depth, independent of sandstone texture. this is probably the result of changes in clay provenance. 4. the th/k ratio does not give an indication of the lithofacies in this study. clay-bearing samples (those of lithofacies ii and iii) have th/k ratios of 3 x 10-4 – 10 x 10-4, whereas in the lithofacies i samples this 364 ratio varies widely (th/k: 2 x 10-4 – 22 x 10-4) depending on the content of feldspar and heavy minerals. 5. thorium and uranium signals are primarily indicative of the presence of heavy minerals in the sediments. no correlation was found between the u content and the total organic carbon content. 6. the volume of individual samples investigated in the laboratory is much smaller than the volume ‘seen’ by the sng logging tool in the borehole. however, most samples were selected from intervals with stable concentrations of th, u and k, so that the samples are inferred to be representative for the depth intervals from which they were taken. therefore, the geochemical results obtained in the laboratory can in general be compared directly to log data. only few of the samples were taken from depth intervals with strongly varying concentrations of th, u and k. 7. a major discrepancy was observed between log data and laboratory gamma spectrometry data. before interpreting the borehole log from west lulu-3, the k concentrations should be multiplied by 0.8, and the th concentrations by 0.95. the u concentrations cannot be corrected. acknowledgements k. carlsen, m. christensen, v. knudsen, h. møller, s. nguyen, a. steffensen and i. søndergaard are thanked for technical assistance. f. larsen (geological survey of denmark and greenland), f. engstrøm (mærsk olie og gas as) and s. talman (technical university of denmark) are thanked for critically reading the manuscript. special thanks are extended to the referees a. hurst (university of aberdeen) and j. andsbjerg (geological survey of denmark and greenland) for valuable advice. references adams, j.a.s. & weaver, c.e. 1958: thorium-to-uranium ratios as indicators of sedimentary processes: example of concept of geochemical facies. american association of petroleum geologists bulletin 42, 387–430. aja, s.u., rosenberg, p.e. & kittrick, j.a. 1991: illite equilibria in solutions: i. phase relationships in the system k2o–al2o3–sio2–h2o. geochimica et cosmochimica acta 55, 1353–1364. andsbjerg, j. 2003: sedimentology and sequence stratigraphy of the bryne and lulu formations, middle jurassic, northern danish central graben. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 301–347 (this volume). berstad, s. & dypvik, h. 1982: sedimentological evolution and natural radioactivity of tertiary sediments from the central north sea. journal of petroleum geology 5, 77–88. bjørkum, p.a. & gjelsvik, n. 1988: an isochemical model for formation of authigenic kaolinite, k-feldspar and illite in sediments. journal of sedimentary petrology 58, 506–511. bjørkum, p.a., walderhaug, o. & aase, n.e. 1993: a model for the effect of illitization on porosity and quartz cementation of sandstones. journal of sedimentary petrology 63, 1089–1091. bjørlykke, k. & egeberg, p.k. 1993: quartz cementation in sedimentary basins. american association of petroleum geologists bulletin 77, 1538–1548. cowan, d.r. & myers, k.j. 1988: surface gamma ray logs: a correlation tool for frontier areas: discussion. american association of petroleum geologists bulletin 72, 634–636. curtis, c.d. 1983: link between aluminium mobility and destruction of secondary porosity. american association of petroleum geologists bulletin 67, 380–393. ehrenberg, s.n. 1990: relationship between diagenesis and reservoir quality in sandstones of the garn formation, haltenbanken, mid-norwegian continental shelf. american association of petroleum geologists bulletin 74, 1538–1558. ehrenberg, s.n., aagaard, p., wilson, m.j., fraser, a.r. & duthie, d.m.l. 1993: depth-dependent transformation of kaolinite to dickite in sandstones of the norwegian continental shelf. clay minerals 28, 325–352. ellis, d.v. 1982: correction of ngt logs for the presence of kcl and barite muds. society of professional well log analysts 23rd annual logging symposium, july 6–9, 1982. transactions, paper o, 12 pp. engstrøm, f. 1981: fortolkningsgrundlag for sng-log, 469 pp. unpublished ph.d. thesis, danmarks tekniske universitet, lyngby, danmark. hansley, p.l. & nuccio, v.f. 1992: upper cretaceous shannon sandstone reservoirs, powder river basin, wyoming: evidence for organic acid diagenesis. american association of petroleum geologists bulletin 76, 781–791. hassan, m., hossin, a. & combaz, a. 1976: fundamentals of the differential gamma ray log – interpretation technique. society of professional well log analysts 17th annual logging symposium, june 9–12, 1976. transactions, paper h, 18 pp. hurst, a. 1990: natural gamma-ray spectroscopy in hydrocarbonbearing sandstones from the norwegian continental shelf. in: hurst, a., lovell, m.a., & morton, a.c. (eds): geological applications of wireline logs. geological society special publication (london) 48, 211–222. hurst, a. & milodowski, a. 1994: characterisation of clays in sandstones: thorium content and spectral log data. society of professional well log analysts 16th european formation evaluation symposium, october 11–13, 1994. transactions, paper s, 18 pp. iaea 1987: preparation and certification of iaea gamma-ray spectrometry reference materials. report file number iaea/rl/148. vienna: iaea (international atomic energy agency). 365 366 imam, m.b. & trewin, n.h. 1991: factors contributing to high gamma-ray levels in upper jurassic sandstone reservoirs of the claymore oilfield, north sea. marine and petroleum geology 8, 452–460. johannessen, p.n. & andsbjerg, j. 1993: middle to late jurassic basin evolution and sandstone reservoir distribution in the danish central trough. in: parker, j.r. (ed.): petroleum geology of northwest europe: proceedings of the 4th conference, 271–283. london: geological society. lindgreen, h. 1985: diagenesis and primary migration in upper jurassic claystone source rocks in north sea. american association of petroleum geologists bulletin 69, 525–536. michelsen, o., nielsen, l.h., johannessen, p.n., andsbjerg, j. & surlyk, f. 2003: jurassic lithostratigraphy and stratigraphic development onshore and offshore denmark. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 147–216 (this volume). milliken, k.l. & land, l.s. 1991: reverse weathering, the carbonate–feldspar system, and porosity evolution during burial of sandstones. american association of petroleum geologists bulletin 75, 636 only (abstract). nielsen, b.l., løvborg, l., sørensen, p. & mose, e. 1987: gammaray analysis for u, th and k on bulk cutting samples from deep wells in the danish subbasin and the north german basin, 82 pp. unpublished report, risø national laboratory, denmark. postma, d. 1983: pyrite and siderite in brackish and freshwater swamp sediments. american journal of science 282, 1151–1183. schlumberger 1982: natural gamma-ray spectrometry. essentials of n.g.s. interpretation, 69 pp. houston, texas: schlumberger educational services. vernik, l. & nur, a. 1992: petrophysical classification of siliciclastics for lithology and porosity prediction from seismic velocities. american association of petroleum geologists bulletin 76, 1295–1309. waples, d.g. 1980: time and temperature in petroleum formation: application of lopatin’s method to petroleum exploration. american association of petroleum geologists bulletin 64, 916–926. manuscript received 26 august 1994; revision accepted 16 october 1997. geological survey of denmark and greenland bulletin 35, 2016, 9-12 9© 2016 geus. geological survey of denmark and greenland bulletin 35, 9–12. open access: www.geus.dk/publications/bull a motorway was constructed in 2010–2016 through the suburbs of the city of silkeborg (fig. 1). the danish road directorate wished to climate-proof the motorway against adverse future climate changes. the directorate collaborated with the geological survey of denmark and greenland (geus) to study the hydrological conditions. studies of historical and projected climate-change-driven variations in groundwater levels in relation to urbanised hydrological fluxes were conducted by kidmose et al. (2013, 2015). during the construction of the motorway, miocene and quaternary deposits were exposed in the slopes of the gudenå valley and late-glacial glaciofluvial deposits were found in the valley floor. this paper focuses on the miocene sediments and their influence on the local hydrological conditions. at silkeborg the gudenå valley is c. 35 m deep (fig. 1). the surrounding terrain is a till plain. in the slope of the valley, glaciofluvial sand is found below the till. miocene deposits are found below the glaciofluvial sand. the floor of the gudenå valley is covered by c. 15 m thick glaciofluvial deposits, which rest on miocene deposits. in borehole no. dgu 87.907 49 m of miocene deposits belonging to the vejle fjord formation are recorded, consisting primarily of marine clay with minor occurrences of sandy deposits. about 12 km south of silkeborg lower miocene deposits are seen in outcrops and boreholes (fig. 2). here the fluvial addit member of the billund formation (rasmussen at al. 2010) is separated from the underlying marine vejle fjord formation by a sharp erosional contact (rasmussen 2014). miocene deposits at silkeborg, jylland, and their influence on hydrology peter roll jakobsen, erik skovbjerg rasmussen, karen dybkjær and jacob kidmose 1 km silkeborg gudenå clayey till sandy till glaciofluvial sand outwash plain peat aeolian sand miocene deposits well no. 87.907 fig. 1. geological map of the silkeborg area. dashed line: motorway. 1 km 0 l ev el ( m ) 50 50 100 100 continental sand and gravel marine clay marine sand coal pre-miocene eastwest addit mb billund fm vejle fjord fm ? 8 7 .9 0 7 fi g. 3 quaternary deposits continental clay fig. 2. east–west profile about 12 km south of silkeborg (from rasmussen 2014). a composite log from silkeborg (fig. 3) and data from well no. dgu 87.907 are shown to the left. planar cross-bedding parallel bedding clay trace fossil sample for palynology hummocky cross-bedding sedimentary structures lithology fig. 4a fig. 4c fig. 4d fig. 4e fig. 4b a q u it an ia n a d d it m b , b ill u n d f m v ej le f jo rd f o rm at io n 0 m 1 m 2 m 3 m 4 m 5 m 6 m 7 m 8 m 9 m sand gravel silt fine medium coarse fig. 3. composite sedimentological log of the temporarily exposed miocene deposits. 1010 sedimentology the section along the motorway comprised 9 m of miocene deposits (fig. 3). the lower part is characterised by cross-stratified medium-grained sand, dipping c. 30° towards the north (fig. 4a). a few trace fossils (skolithos?) are seen. the cross-stratified sand is sharply overlain by wave-formed, coarse-grained ripples. the crests of the ripples strike se–nw, and crest-to-crest spacing is in the range of 250 cm with amplitudes up to 35 cm (fig. 4b). the ripples show tangential cross-bedding towards the sw. in a nearby exposure, tidal bundles form the base of the section. the presence of clay layers varies systematically and is commonly characterised by double clay layers (fig. 4a). dips of cross-bedding are both sw and ne. the coarse-grained ripples are in turn overlain by a dark brown mud. the mud is succeeded by silt and fine-grained sand, c. 1.5 m thick. hummocky cross-stratifications (hcs) are common, especially in the upper part of the section. these are superimposed by 3 m of dark brown mud (fig. 4c) showing a slight increase in grain size upwards where hummocky cross-stratified sands are common (fig. 4d). a sharp boundary separates the mud from an overlying 2 m thick section of mediumto coarse-grained sand and gravel. this coarse-grained section is composed of tabular co-sets of cross-stratified beds dipping towards the south (fig. 4e). fig. 4. details of the miocene deposits. the stratigraphic positions of the photos are indicated on fig. 3. a: cross-stratified medium grained sand. double clay layers are indicated with small arrows. trace fossil is indicated with larger arrow. b: wave-formed, coarse grained ripples. c: thick dark mud succession. d: hummocky cross-stratification (hcs). e: tabular co-sets of cross-stratified beds. b c d e a 11 bioand chronostratigraphy and depositional environment in order to confirm the miocene age of the described succession and to achieve a more precise dating, two sediment samples were selected for palynological analysis. the stratigraphic positions of the samples are shown in fig. 3. one of the samples was almost barren, while the other contained a rich assemblage of organic particles dominated by bisaccate and non-saccate pollen. in addition, the sample contained a moderately rich and diverse dinoflagellate cyst (dinocyst) assemblage together with a few wood particles, cuticle, acritarchs and freshwater algae. this assemblage indicates a marine, inner neritic depositional environment with a high influx of freshwater (tyson 1995). the dinocyst assemblage is dominated by two species of the genus homotryblium: h.? additense (fig. 5a) and h. plectilum. among several other dinocyst taxa, a single specimen of the stratigraphically important species chiropteridium galea was found (fig. 5b). the dinocyst assemblage refers the sample to the chiropteridium galea zone (dybkjær & piasecki 2010). this dinocyst zone is dated to the early aquitanian (earliest miocene) and the age of the sample is 23.03–22.36 ma. palaeogeography the sand and gravel in the lower part of the section were deposited during an overall regression of the billund formation in the early miocene (rasmussen et al. 2010). the gravel was probably originally deposited in a fluvial environment during the most extended regression. the crossstratified sand in the lower part was formed in a marine bar that migrated landwards. the tidal bundles were formed by both ebb and flood currents, as indicated by the bipolar dips of cross-bedding, in an adjacent tidal inlet. the overlaying wave-formed, coarse-grained ripples were formed by marine reworking (leckie 1988) of the coarse-grained fluvial sediments laid down during maximum regression and now forms a transgressive lag (plint 1988). the depositional water depth of the coarse-grained ripples may lie in the range of 15 to 60 m (leckie 1988) – most likely in the lower end as the sea-level changes during this part of the miocene was c. 25 m (miller et al. 2005). the strike of the crests of the coarse-grained rippels, se–nw, indicates the trend of the palaeo-shoreline (leckie 1988). the succeeding mud and hcs-dominated silt and fine-grained sand were deposited in slightly deeper water, in the offshore transition zone. the mud-dominated part with few intercalations of hcs’s was deposited offshore near the storm wave base. the assemblage of organic particles indicates that the sediment was deposited in a marine depositional setting near the coast. the two homotryblium species further indicate that the palaeoenvironment was marine but probably with lowered salinity (dybkjær 2004). these interpretations support the sedimentological interpretations and palaeogeographic maps for the earliest miocene of jylland, indicating that large river and delta systems existed, which transported large amounts of freshwater and sediment from the north to the middle part of jylland (rasmussen et al. 2010). 20 μm20 μm a b fig. 5. a: homotryblium? additense. b: chiropteridium galea. fig. 6. palaeogeographical reconstructions of the silkeborg area. a: early aquitanian (earliest miocene, vejle fjord formation) tidal-dominated marine-barrier system. b: late aquitanian (addit member, billund formation) fluvial environment. grey line: motorway. a b1 km 1212 authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: prj@geus.dk the coarse-grained sand and gravel at the top of the studied succession, that sharply overlie the marine deposits, were deposited in a fluvial environment, the addit member of the billund formation (rasmussen et al. 2010; fig. 6). the dramatic change in the depositional environment was partly a result of an eustatic sea-level fall and partly a result of inversion of the norwegian–danish basin (rasmussen 2014). the latter resulted in marked incision in the middle and northern part of jylland during the late aquitanian. the miocene succession in silkeborg shows a strong resemblance to successions in boreholes and exposures about 12 km south of silkeborg (fig. 2). hydrology on the floor of the gudenå valley, wells with screens in the miocene deposits have artesian hydraulic heads, whereas the hydraulic head in the overlying glaciofluvial sediments is in hydraulic contact with the gudenå. this shows that the alternating miocene layers of the vejle fjord formation form a hydraulic barrier between deeper groundwater and the surficial glaciofluvial aquifer that is in contact with the motorway (fig. 7). in the higher terrain the measured hydraulic head in the quaternary glaciofluvial sand is very different from the hydraulic head measured in the underlying miocene deposits. this is because the 3 m thick miocene mud unit (fig. 4c) acts as a barrier. the addit member, however, is in hydraulic contact with the glaciofluvial sand (fig. 7). the hydraulic connection between the glaciofluvial sand found in the slopes of the valley and the glaciofluvial deposits in the valley floor is also affected by the miocene deposits as the 3 m thick mud unit separates them. the hydrogeological relations between the miocene deposits and the quaternary deposits illustrate the importance of applying detailed field-site geological evidence to get an impression of the local groundwater flow. acknowledgement the danish road directorate is thanked for access to the field site and funding. references dybkjær, k. 2004: morphological and abundance variations in homotryblium-cyst assemblages related to depositional environments; uppermost oligocene – lower miocene, jylland, denmark. palaeogeography, palaeoclimatology, palaeoecology 206, 41–58. dybkjær, k. & piasecki, s. 2010: neogene dinocyst zonation in the eastern north sea basin, denmark. review of palaeobotany and palynology 161, 1–29. kidmose, j., refsgaard, j.c., troldborg, l., seaby, l.p. & escrivà, m.m. 2013: climate change impact on groundwater levels: ensemble modelling of extreme values. hydrology and earth system sciences 17, 1619–1634. kidmose, j., troldborg, l., refsgaard j.c. & bischoff, n. 2015: coupling of a distributed hydrological model with an urban storm water model for impact analysis of forced infiltration. journal of hydrology 525, 506–520. leckie, d. 1988: wave-formed, coarse-grained ripples and their relationship to hummocky cross-stratification. journal of sedimentary research 58, 607–622. miller, k.g. et al. 2005: the phanerozoic record of sea-level changes. science 310, 1293–1298. plint, a.g. 1988: sharp-based shoreface sequences and “offshore bars” in the cardium formation of alberta; their relationship to relative changes in sea level: in: wilgus, c.k. et al. (eds): sea-level changes: an integrated approach: sepm, special publication 42, 357–371. rasmussen, e.s. 2014: development of an incised-valley fill under the influence of tectonism and glacio-eustatic sea-level change: valley morphology, fluvial style and lithology. journal of sedimentary research 84, 278–300. rasmussen, e.s., dybkjær, k. & piasecki, s. 2010: lithostratigraphy of the upper oligocene – miocene succession in denmark. geological survey of denmark and greenland bulletin 22, 92 pp. tyson, r. 1995: sedimentary organic matter: organic facies and palynofacies, 615 pp. london: chapman & hall. clayey till meltwater sand diamicton meltwater plain sand clay quaternary deposits miocene deposits w el l n o . 8 7 .9 0 7 addit member vejle fjord formation sand sand/clay level (m) 60 50 40 30 20 10 0 –10 –20 fig. 7. conceptual geological model along the motorway alignment. geological survey of denmark and greenland bulletin 31, 2014, 55-58 55 estimating thermal conductivity from lithological descriptions – a new web-based tool for planning of ground-source heating and cooling claus ditlefsen, inga sørensen, morten slott and martin hansen it is the overall policy of the danish government that by 2050 electricity, heating and transport will be 100% based on renewable energy. in order to reach this goal a number of different green technologies will have to interact. in areas with no district heating, ground-source heating by heat pump technology (sanner 2011) could well be one of the solutions. the potential energy extraction from closed-loop boreholes for ground-source heating depends to a large degree on the thermal conductivity of the surrounding geological formations, although other parameters such as the thermal gradient and the extent of groundwater flow also affect the transport of heat to the borehole. initial estimates indicate that in denmark there may be as much as 40% difference between the most and the least favourable geological conditions, determined by the thermal conductivity of the different sediment or rock types alone (vangkilde-pedersen et al. 2012). therefore specific knowledge of the thermal conductivity of the geological formations is essential when estimating the optimal drilling depth and the number of boreholes required for a specific plant. in co-operation with research and industrial partners, the geological survey of denmark and greenland is conducting a three-year project with the title ‘geoenergy, tools for ground-source heating and cooling based on closed-loop boreholes’ (www.geoenergi.org). the objective of the project is to acquire knowledge and develop tools and best practices for the planning, design and installation of shallow geothermal energy systems. this paper describes a web-based tool developed to estimate the thermal conductivity in the area surrounding a potential new plant. the tool was developed within the geoenergy project and can be used by administrators, energy planners and drillers of closed-loop boreholes. thermal conductivity of shallow danish sediments the thermal conductivity of sediments or rocks depends on their mineral composition, the texture, and the water content. above the water table, where air is present in the pore spaces, sediments generally have a low thermal conductivity. hence information about the position of the water table in the borehole is important when planning a new site. relatively few investigations of thermal properties of danish sediments have been carried out (balling et al. 1981; porsvig 1986) and thermal conductivity values published in international literature show broad ranges for the individual sediment types (e.g. banks 2008; vdi 2010). this is generally the case for clayey sediments and particularly for glacial till, see vangkilde-pedersen et al. (2012) for details. therefore a programme was initiated to investigate the thermal properties of common shallow sediments. the work focused on determining the thermal conductivity of prevalent watersaturated sediments. the range of the thermal conductivities within common sediment types was determined from measurements of 51 samples from well-characterised exposures at different localities (figs 1, 2). the samples were water saturated in the laboratory and placed in a thermal cupboard at 20°c for at least 16 hours before measurements were made. the thermal conductivity was determined using the needle probe method (wechsler 1992; hukseflux 2003). detailed © 2014 geus. geological survey of denmark and greenland bulletin 31, 55–58. open access: www.geus.dk/publications/bull 10°e 14°e 55° 57°n 50 km bornholm denmark sweden fig. 3 germany fig. 1. map showing sample locations. one sample from northern germany was kindly provided by reinhard  kirsch, landesamt  schleswigholstein. for sample details see ditlefsen & sørensen (2014). the arrow shows the location of the map included in fig. 3. http://www.geoenergi.org 5656 descriptions of sampling and laboratory procedures, as well as data analysis and statistical analysis were published by ditlefsen & sørensen (2014) and sørensen et al. (in press). a summary of the results is provided in table 1. each sample has been measured 2 to 5 times and an average value representing the sample was calculated. the variation amongst samples of the same sediment type is given as one standard deviation using the average values of each sample. the results indicate that the different sediment types have thermal conductivities within characteristic ranges with one standard deviation corresponding to approximately 20% of the mean. this further implies that it will be possible to estimate the thermal conductivity around a specific borehole from thorough descriptions of borehole samples alone. the national borehole database jupiter the geological survey of denmark and greenland has acquired data on boreholes since 1926 in accordance with the danish water supply legislation. the data include information about location, construction, geology, water table and groundwater chemistry (hansen & pjetursson 2011). samples from approximately one third of the boreholes have been described and interpreted by geologists; the rest have been described by drillers in the field. since 1969, drillers have also been obliged to submit representative borehole samples to the survey where they are described according to strict standards as outlined by larsen et al. (1995). in addition, geological interpretations of age and depositional environment are made (gravesen & fredericia 1984). the data are stored in the national borehole database, jupiter, which can be accessed on the internet free of charge (www.geus.dk). the database includes data on more than 270 000 boreholes, corresponding to about six boreholes per square kilometre. in 2001, training and certification of drillers operating in denmark became mandatory, including instruction in making a simple but rigorous description of borehole samples. this allows for an overall assessment of the character and possible origin of samples that have been described in the field by drillers. all in all, the national borehole database provides planners, drillers and administrators with a unique possibility to evaluate local geological conditions at a given site. estimating thermal conductivity values from sediment descriptions as described above, it has been possible to establish a relationship between lithology and thermal conductivity for a number of common danish sediment types. the national database holds a large number of lithological descriptions from throughout the country, and by combining the lithological and thermal conductivity data the borehole database can be used in a new way. to do this it has been necessary to develop a routine that could relate a lithological description to one of the sediment groups in table 1. this task was facilitated by the structure of the lithological table in jupiter, where different components of the lithological description e.g. rock type, minor components, mineralogy, grain size, overall interpretation etc. are stored with unique codes a b fig. 2. a: sampling of clayey till. b: a sediment sample with a needle probe installed. gyttja 3 0.68 0.58–0.86 0.15 smectite-rich clay 3 0.98 0.80–1.14 0.17 silty clay 10 1.15 0.90–1.42 0.17 chalk* 4 1.62 1.49–1.80 0.13 mica-rich, fine-grained sand 8 1.81 1.48–2.19 0.27 till 19 1.89 1.40–2.66 0.30 glacial sand, gravelly 4 2.24 1.98–2.43 0.19 pure quartz sand 3 2.75 2.41–3.34 0.51 table 1. thermal conductivity of some common, shallow, danish sediments * selected data from balling et al. (1981). measurements were conducted with a needle probe (hukseflux 2008) using water-saturated samples. sediment type number of samples average thermal conductivity range one standard deviation w mk–1 w mk–1 w mk–1 57 in individual data fields, which makes rigorous queries into the lithological data possible. in this way most samples described and interpreted by geologists could fairly easily be assigned to one of the sediment groups in table 1. in addition granite, gneiss and sedimentary rocks found near the surface on the island of bornholm (fig. 1) are tentatively ascribed thermal conductivities in accordance with vdi (2010). water-lain sediments consisting of alternating layers of sand and clay are tentatively ascribed an average thermal conductivity of 1.5 w mk–1 in accordance with the values for sand and clay (table 1). for samples where only the overall sediment type was noted by the driller, interpretations had to be made (table 2). finally, to compensate for the fact that sediments which are not water saturated have reduced thermal conductivities (vdi 2010), all deposits above the water table in the borehole are tentatively ascribed a conductivity of 1.0 w mk–1. table 2. interpretation of thermal conductivity * sediment type according to driller. § the values are based on measurements of thermal conductivity (table 1) and our interpretations of sediment types. $ clay deposits with sand, gravel or stones are interpreted as till. other clay deposits are interpreted as clay deposited in water. sediment type* supplementary information from driller suggested thermal conductivity (w mk–1)§ sand 2.24 limestone 1.62 clay 1.15 clay$ containing, sand, gravel or stones 1.89 clay, sand, and stones$ 1.89 fig. 3. results from the web application shown in a standardised report window. http://geuskort.geus.dk/termiskejordarter/ –2 print 1.47 1.20 1.95 2.18 alternating sand and clay d ep th (m ) water table n um be r o f b or eh ol es estimated thermal conductivity (w mk–1) 2 2 1 1 1 sediment distribution (%) 0 25 50 75 100 125 1.65 1 2 3 1000 glacial till mica-rich sand smectite-rich clay chalk and limestone gyttja and peat dry sediment above water table 1.00 1.81 1.50 0.98 1.62 0.68 1.15 1.89 silty clay granite, gneiss and sandstone quartz sand shale sand and gravel 2.75 2.20 2.24 3.00 89.257 89.90 56°8.7´n 10°8.1´e 100 m http://geuskort.geus.dk/termiskejordarter/ 5858 it is the intention of the web application to show the expected thermal conductivity in a new project area from existing lithological descriptions and further to show the variation in thermal conductivity with depth as a function of the lithological variations. the calculations include all available boreholes within a user-defined rectangle, and the estimated thermal conductivity is calculated in depth intervals of 25 m from the available lithological descriptions. furthermore, it is required that at least 80% of the interval is covered by descriptions that can be related to a thermal conductivity value. if this requirement is not met, the borehole is excluded from the calculations for the specific depth interval. within each interval the percentage of each sediment type is then calculated and from this distribution the resulting thermal conductivity of the interval is estimated from the reference values for water-saturated sediments or rocks (fig. 3). furthermore, the average depth to the water table is calculated from the most recent soundings in each borehole. above this depth a reduced thermal conductivity of 1 w mk–1 is estimated overruling the thermal conductivity related to the water saturated sediment. the web application is available at http://geuskort.geus. dk/termiskejordarter/. from the initial map, the user can zoom in on the relevant project area and see all boreholes. by using the box search button and clicking on the individual boreholes, a standard lithological report appears and the quality of the individual borehole data can be accessed. by dragging a rectangle over one or more boreholes, calculations based on the selected boreholes are made as described above. in order to obtain a reliable estimate of the thermal conductivity, it is important that the selected boreholes represent the geology at the new site, and a rectangle size of not more than 1 km2 is recommend unless an initial data inspection indicates otherwise. the results are shown in a standardised report window (fig. 3). the report shows the average lithological composition of each 25 m interval as bars and a plot of expected thermal conductivity values versus depth. a plot of the depth to the water table calculated from soundings in the area is shown to the left. the report also contains a locality map that shows the boreholes in the area and a legend that includes average thermal conductivities of different sediment and rock groups. the properties of the different groups can easily be adjusted or more groups can be added by the web administrator when more information about the thermal conductivity of different sediments and rocks becomes available. so far, the web application has been tested and released. the next step will be to introduce it to different end users such as administrators, drillers and energy planners. we also plan to conduct a number of interviews to get feedback, which may lead to adjustment of the system. acknowledgement the eudp programme of the danish energy agency is thanked for financial support of the geoenergy project. references balling, n., kristiansen, j.i., breiner, n., poulsen, k.d., rasmussen, r. & saxov, s. 1981: geothermal measurements and subsurface temperature modelling in denmark. geoskrifter 16, 176 pp. banks, d. 2008: an introduction to thermogeology: ground source heating and cooling, 351 pp. oxford: blackwell publishing. ditlefsen, c. & sørensen, i. 2014: d6 overfladenære jordarters termiske egenskaber (in danish with english summary), 30 pp. http://geoenergi. org/xpdf/d6_ jordarters_termiske_egenskaber.pdf gravesen, p. & fredericia, j. (eds) 1984: zeus-geodatabase system. borearkivet. databeskrivelse, kodesystem og sideregistre. danmarks geologiske undersøgelse serie d 3, 259 pp. hansen, m. & pjetursson, b. 2011: free, online danish shallow geological data. geological survey of denmark and greenland bulletin 23, 53–56. hukseflux 2003: tp02 non-steady-state probe for thermal conductivity measurement. tp02 manual v1209. delft: hukseflux. hukseflux 2008: tpsys02  thermal conductivity measurement system user manual tpsys02 manual v0806.doc. delft: hukseflux. larsen, g., frederiksen, j., willumsen, a., fredericia, j., gravesen, p., foged, n., knudsen, b. & baumann, j. 1995: a guide to engineering geological soil description. danish geotechnical society bulletin 1e, 130 pp. porsvig, m. 1986: varmeovergangsforhold omkring jordslanger. energiministeriets varmepumpeforskningsprogram 33, 56 pp. sanner, b. 2011: concept and feasibility studies. in: mc corry, m. & jones, g.l. (eds): geotrainet training manual for designers of shallow geothermal systems, 21–24. brussels: geotrainet. sørensen, i., palola, m.a. & ditlefsen, c. 2014: d8 guidelines for equipment, methods and calibration, part 1: measurement of thermal conductivity. geoenergy report, 37 pp. www.geoenergy.dk vangkilde-pedersen, t., ditlefsen, c. & højberg, a.l. 2012: shallow geothermal energy in denmark. geological survey of denmark and greenland bulletin 26, 37–40. vdi (verein deutscher ingenieure) 2010: thermische nutzung des untergrundes: grundlagen, genehmigungen, umweltaspekte. richtlinie 4640, blatt 1, 33 pp. düsseldorf: verein deutscher ingenieure. wechsler, a.e. 1992: the probe method for measurement of thermal conductivity. in: maglic, k. d., cezairliyan, a. & peletsky, v.e. (eds): compendium of thermophysical property measurement methods 2, recommended measurement techniques and practices, 161–185. new york: plenum press. authors’ addresses c.d., geological survey of denmark and greenland, lyseng allé 1, dk-8270, højbjerg, denmark. e-mail: cd@geus.dk m.s. & m.h., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. i.s., via university college, campus horsens, chr. m. østergaards vej 4, dk-8700 horsens, denmark. http://geuskort.geus.dk/termiskejordarter/ http://geuskort.geus.dk/termiskejordarter/ http://geoenergi.org/xpdf/d6_jordarters_termiske_egenskaber.pdf http://geoenergi.org/xpdf/d6_jordarters_termiske_egenskaber.pdf http://www.geoenergy.dk geological survey of denmark and greenland bulletin 38, 2017, 37-40 37 the lower palaeozoic succession in scandinavia includes several excellent marine source rocks notably the alum shale, the dicellograptus shale and the rastrites shale that have been targets for shale gas exploration since 2008. we here report on samples of these source rocks from cored shallow scientific wells in southern sweden. the samples contain both free and sorbed hydrocarbon gases with concentrations significantly above the background gas level. the gases consist of a mixture of thermogenic and bacterially derived gas. the latter likely derives from both carbonate reduction and methyl fermentation processes. the presence of both thermogenic and biogenic gas in the lower palaeozoic shales is in agreement with results from past and present exploration activities; thermogenic gas is a target in deeply buried, gas-mature shales in southernmost sweden, denmark and northern poland, whereas biogenic gas is a target in shallow, immature-marginally mature shales in south central sweden. we here document that biogenic gas signatures are present also in gas-mature shallow buried shales in skåne in southernmost sweden. in south central sweden (västergötland, östergötland, närke and öland, fig. 1), shallow (present burial <150 m) immature to marginally mature bituminous shale has been known for decades to contain gas and is currently under exploration (see summary in schultz et al. 2015). since 2009, many deep (>800 m) exploration drillings in northern poland (including the lebork s-1 well, lehr & keeley 2016), the vendsyssel-1 well in denmark (ferrand et al. 2016) and the a3-1, b2-1, c4-1 wells in skåne in southernmost sweden (pool et al. 2012), have demonstrated that the lower palaeozoic shale succession contains gas also in these areas (fig. 1). in denmark and skåne, the average gas content is 30 ft3 gas per ton of rock in the organicrich alum shale formation (pool et al. 2012; ferrand et al. 2016). the equivalent shale formation in northern poland contains up to 268 ft3 gas per ton of rock in the lebork s-1 well (lehr & keeley 2016), which is comparable to the content within the core area of north american shale gasproducing formations (e.g. jarvie 2012). in this study, we present empirical data on the composition and isotope signatures of gas measured in shallow (<158 m) core samples from five scientific core drillings in southern sweden, viz. albjära-1, lönstorp-1, gislövshammar-2, hällekis-1 and djupvik-1 (fig. 1, table 1). sampling and analyses were performed in 1991–1992 as part of the energy research project efp-1313/88-2 and the pre-westphalian source-rocks in northwest europe (prewsor) project (schovsbo & laier 2012). this paper aims at further characterising the gas composition. samples and analyses the molecular composition and isotope signatures of the occurring gases were measured in 27 core samples (table 1). the samples were selected during drilling and consist of 4–8 cm long core intervals with a diameter of 5.5 cm. the samples from albjära-1 and lönstorp-1 were sealed in metal containers and stored at –18°c until they reached the laboratory for analysis. the free gas was subsequently anageneration and origin of natural gas in lower palaeozoic shales from southern sweden niels hemmingsen schovsbo and arne thorshøj nielsen djupvik-1 albjära-1 a3-1 c4-1 gislövshammar-2 lebork s-1 hällekis-1 b2-1 vendsyssel-1 lönstorp-1 figur 1. one collumn wide öland baltic sea bornholm sweden denmark skåne västergötland östergötland närke poland lower palaeozoic strata caledonian front exploration well scientific borehole 200 km fig. 1: location of wells mentioned and occurrence of lower palaeozoic strata in southern scandinavia. modified from nielsen & schovsbo (2015). © 2017 geus. geological survey of denmark and greenland bulletin 38, 37–40. open access: www.geus.dk/publications/bull 3838 lysed by the geological survey of denmark and greenland (geus) by puncturing the containers through a septum before opening. samples from gislövshammar-2, hällekis-1 and djupvik-1 were analysed by the federal institute for geosciences and natural resources (bgr), germany. these samples were stored at –18°c at the drill site and subsequently transferred to a container filled with liquid nitrogen at –196°c. the free gas was measured by allowing the deep-frozen sample to equilibrate to room temperature in a sealed container. for all samples the sorbed gas in the rock matrix was liberated by treating the sample with phosphoric acid following the procedure outlined by faber & stahl (1983). gas composition the sorbed methane concentrations range from 118 to 23867 ppb (micrograms per kg of rock) and the free gas contents from 3 to 13420 ppm (fig. 2). the concentrations thus far exceed the level of background gas of 20–50 ppb for methane as defined by whiticar (1994). the gas content is strongly related to thermal rank of the sampled shales. the gas-mature samples from gislövshammar-2, albjära-1 and lönstorp-1 have c. ten times higher yields than the thermally immature samples from the hällekis-1 and djupvik-1 wells (fig. 2). the gas content appears not to be related to the total organic content in the mature samples. this is exemplified by the fact that the highest yields of sorbed gas are found in the rastrites shale (toc average <1%) and the lowest yields in the alum shale, which on average contains 9% toc in skåne and on bornholm (fig. 2). isotope composition and carbon isotope signatures the gas molecule and isotopic compositions of the analysed samples plot within the bacterial to thermogenic fields in a bernard diagram (fig. 3a ). the microbial gas is characterised by much higher negative isotope values than seen in the thermogenic-sourced methane. this signature is most clearly expressed in the free gas samples from mature shales that also have 10 to 100 times higher c1/(c2+c3) molecular ratios than the sorbed gas, which is also typical for biogenic gas (fig. 3a). one column wide alme. ra str ite s linde. 480 e. c . m a m .c am br ian 500 510 470 460 450 440 430 490 sil ur ian fu ro ng ian ea rly o rd ov ici an m .o rd ov ici an la te o rd ov ici an grey shalelimestone siltstone black shale -80 -60 -40 δ13cmethane ‰ pdbtoc wt% 0 5 1510 figure 2. ba al um s ha le g dicello. tø ye n sh ale k b gas: sorbed free mature immature 6 000 ppb 2 000 ppm 13 000 ppm 25 000 ppb fig. 2. stratigraphy of the lower palaeozoic shales. a: total organic carbon (toc) content in shales from the skåne–bornholm area (modified from schovsbo 2003). b: methane isotope composition of sorbed and free gas from samples of thermally mature shale in the albjära-1, gislövshammar-2, and lönstorp-1 wells and from samples of thermally immature shale in the hällekis-1 and djupvik-1 wells. e.c.: early cambrian. g: gislöv fm. b: bjørkåsholmen fm. k: komstad limestone. alme.: almelund shale. dicello.: dicellograptus shale. linde.: lindegård formation. green area in b outlines the variation field defined by samples from the djupvik-1 and hällekis-1 wells. table 1. wells and analysed samples§ well formation n depth range (m) gislövshammar-2 tøyen s. 1 19.8 gislövshammar-2 alum s. 4 31.8–88.3 djupvik-1 alum s: 1 2.0 hällekis-1 tøyen s. 2 9.9–17.3 hällekis-1 alum s, 4 23.4–39.8 albjära-1 almelund 1 99.6 albjära-1 tøyen s. 3 114.8–134.5 albjära-1 alum s. 5 139.6–157.0 §full analytical results are available on request from the first author. 39 the propane versus ethane isotope gas signature is suggested to reflect the maturity of the source and this relationship can be used for a gas-to-source-rock correlation (whiticar 1994). the analysed gas signatures follow this prediction, as immature alum shale samples from the hällekis-1 well plot with relatively depleted isotope signatures of propane and ethane compared to the thermally mature alum shale samples (fig. 3b). isotope data from lithuanian and polish oil and gas reservoirs in middle cambrian sandstone, sourced by alum shale (kotarba & lewan 2013), plot with intermediate signatures (fig. 3b). the thermally mature samples (gislövshammar-2 well), however, exhibit a considerable variation in propane isotope composition from –22 to –39‰ pdb, suggesting a mixture of differently derived gases. this may be caused by addition of bacterially derived ethane and/or biodegradation of a propane component (cf. whiticar 1994). the thermally immature alum shale samples from the djupvik-1 well (marked with 0.32 %gr in fig. 3b) and samples from tøyen shale from the hällekis-1 well (marked as tøyen 0.53%gr in fig. 3b) plot within an apparent gas maturity of 0.8–1.2% ro according to the values indicated in fig. 3b (whiticar trend line), i.e. with much higher maturities than measured, suggesting that gas migration has occurred. in the hällekis-1 well the migrated gas may have formed in response to intrusion of permo-carboniferous dikes that locally matured the shales in south central sweden (cf. schultz et al. 2015). in the vicinity of the djupvik-1 drill site-mature shale and igneous activity are unknown and the relatively high maturity remains puzzling. deuterium isotopes measurements of deuterium and carbon isotopes in methane offer additional information on the source of the natural gases and on the processes that may have modified their composition (whiticar 1994). figure 3c shows the deuterium versus carbon isotope composition of methane in the analysed samples. the gas composition exhibits a % ro 0.1 1 10 100 1000 thermogenic bacterial ke ro ge n ty pe ii typ e i ii microbial oxidation c 1 / ( c 2 + c 3) δ13cmethane (‰ pdb) -80 a -70 -60 -50 -40 -30 3.0 2.0 % ro 1.5 0.7 0.5 alum 0.53 % gr tøyen 0.53 % gr 1.7 % gr poland–lithuania 0.6–1.4 % gr 0.32 % gr 1.1 -20-30-40-50 -40 -36 -32 -28 -24 -20 δ13 c pr op an e (‰ p d b) δ13cethane (‰ pdb) b bacterial carbonate reduction geothermal early mature thermogenic mix and transitionbacterial methyl-type fermentation δ13 c m et ha ne (‰ p d b) δdmethane (‰ smow) -350 -300 -250 -200 -150 -100 -80 -60 -40 -20 c whiticar (1994) general trend line mature offshore poland–lithuania sourced by alum shale (kotorba & lewan 2013) immature fig. 3. compostion and isotope signatures of organic carbon. a: ‘bernard diagram’ showing the ratio c1/(c2+c3) versus δ13cmethane for all samples. this type of diagram is used to determine the origin of the gases and is modified from whiticar (1994). for legend, see fig. 2. b: relationship between δ13cethane and δ13cpropane. tøyen and alum 0.53% gr denote stratigraphy and maturity of samples from the hällekis-1 well. the graptolite reflectance (% gr) is from pedersen et al. (2013); grönvik locality is used for djupvik-1; estimate for offshore poland. c: δd and δ13c for methane. for legend, see fig. 2. the isotope signatures of the various sources are from whiticar (1994). 4040 large degree of scatter, but free gas from thermally mature samples plots in the bacterial carbonate reduction field, whereas free gas from immature samples plots towards the bacterial methyl-type fermentation fields, suggesting that different processes generated the depleted methane isotope compositions (fig. 3c). the sorbed gas compositions in general plot away from bacterial sources, indicating that a thermogenic signature is preserved (fig. 3c). implications for shale-gas prospectivity the gas-isotope composition suggests that significant postgeneration modification occurred although the timing is unknown. the biogenic isotope signature of the gases in skåne resembles similar signatures seen in östergötland (schultz et al. 2015). here schultz et al. inferred that methyl-fermenting processes contributed to the methane content. however, the gas-isotope signature was not as depleted as seen in this study, possibly owing to the mixed shale oil – biogenic nature of the östergötland alum shale play. according to schultz et al. (2015) the biogenic gas was generated after the pleistocene glaciation, as modern meteoric water was able to infiltrate the shale and create the right conditions for bacterial activities. we envisage that similar conditions may have affected the shallowly buried shales in skåne. krüger et al. (2014), however, show that highly mature kerogen has a much smaller microbial generative gas potential than immature to marginally mature kerogen, since thermal maturity limits the amount of easily biodegradable organic matter that can be transformed to methane. conclusions lower palaeozoic shales in south central sweden and southernmost sweden contain natural gas that exceeds the level of background gas. the gas content is strongly related to the thermal rank of the sampled shales, and mature samples have approximately ten times higher yields than the immature samples. the gas is generated by both thermogenic and bacterial processes. the microbial gas signature is most clearly expressed in the free gas samples from mature shales that also have 10 to 100 times higher molecule c1/(c2+c3) ratios than the sorbed gas. migration may have occurred related to gas formation in response to intrusion of permo-carboniferous dikes that locally matured the shales in south central sweden. acknowledgements we thank troels laier (geus) for comments and suggestions to an earlier version of the manuscript. the authors wish to thank mikael erlström and maciej kotarba for constructive comments that improved the paper. references faber, e. & stahl, w. 1983: analytic procedure and results of an isotope geochemical surface survey in an area of the british north sea. geological society special publications (london) 12, 51–63. ferrand, j., demars, c. & allache, f. 2016: denmark – l1/10 licence relinquishment recommendations report. total e&p, memo 1-9. available from: http://www.ft.dk/samling/20151/almdel/efk/bilag/353/1651289.pdf. verified 7.4.2017. jarvie, d.m. 2012: shale resource systems for oil and gas: part 1 – shalegas resource systems. aapg memoir 97, 69–87. kotarba, m.j. & lewan, m.d. 2013: sources of natural gases in middle cambrian reservoirs in polish and lithuanian baltic basin as determined by stable isotopes and hydrous pyrolysis of lower palaeozoic source rocks. chemical geology 345, 62–76. krüger, m., van berk, w., arning, e.t., jiménez, n., schovsbo, n.h., straaten, n. & schulz, h.-m. 2014: the biogenic methane potential of european gas shale analogues: results from incubation experiments and thermodynamic modelling. international journal of coal geology 136, 59–74. lehr, j.h. & keeley, j. 2016: alternative energy and shale gas encyclopedia. 912 pp. john wiley & sons. nielsen, a.t. & schovsbo, n.h. 2015: the regressive early – mid cambrian ‘hawke bay event’ in baltoscandia: epeirogenic uplift in concert with eustasy. earth science reviews 151, 288–350. petersen, h.i., schovsbo, n.h. & nielsen, a.t. 2013: reflectance measurements of zooclasts and solid bitumen in lower palaeozoic shales, southern scandinavia: correlation to vitrinite reflectance. international journal of coal petrology 114, 1–18. pool, w., geluk, m., abels, j. & tiley, g. 2012: assessment of an unusual european shale gas play – the cambro–ordovician alum shale, southern sweden: proceedings of the society of petroleum engineers/european association of geoscientists and engineers unconventional resources conference, vienna, austria, 20–22 march, 2012, 152339. schovsbo, n.h. 2003: the geochemistry of lower palaeozoic sediments deposited on the margins of baltica. bulletin of the geological society of denmark 50, 11–27. schovsbo, n.h. & laier, t. 2012: composition and gas isotope signature of shale samples from 5 scientific wells in sweden. geological survey of denmark and greenland report 2012/17, 1–25. schulz, h.-m., biermann, s., van berk, w., krüger, m., straaten, n., bechtel, a., wirth, r., lüders, v., schovsbo, n.h. & crabtree, s. 2015: from shale oil to biogenic shale gas: retracing organic-inorganic interactions in the alum shale (furongian-lower ordovician) in southern sweden. aapg bulletin 99, 927–956. whiticar, m.j. 1994: correlation of natural gases with their sources. aapg memoir 60, 261–283. authors’ addresses n.h.s., geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350 copenhagen k, denmark; e-mail: nsc@geus.dk a.t.n., department of geosciences and natural resource management, university of copenhagen. øster voldgade 10, dk-1350 copenhagen k, dk. geological survey of denmark and greenland bulletin 33, 2015, 9-12 9 the lower palaeozoic now fully cored and logged on bornholm, denmark niels h. schovsbo, arne t. nielsen and kurt klitten a 558 m long, complete section of the lower palaeozoic succession preserved onshore southern bornholm has been compiled from fi ve fully cored scientifi c wells, carried out between 2005 and 2012. th e scientifi c programme included coring and geophysical logging of the fi ve scientifi c wells that yielded a total of c. 750 m of partially overlapping cores as well as re-logging of water wells and acquisition of shallow seismic data. th e last well drilled, the sommerodde-1, cored the youngest preserved silurian strata on bornholm including strata not exposed in outcrops. th e well penetrated 168.1 m of silurian shales, 42.7 m of upper ordovician shales and 27.9 m of alum shale before it terminated at a depth of 250.3 m in the lower cambrian norretorp member of the læså formation. th e sommerodde-1 well documents that the lower silurian cyrtograptus shale is at least 91.7 m thick and that the rastrites shale is 76.4 m thick. th e complete lower cambrian succession has previously been covered by the 316.0 m deep borggård-1 well that terminated in basement rocks (nielsen et al. 2006). the lower palaeozoic in denmark th e lower palaeozoic is very thick in denmark, locally more than 4 km thick, and the deposition marks the transition from passive margin to a rapidly subsiding foreland basin developed in front of the caledonian orogen towards the south and west (lassen & th ybo 2012). th e palaeozoic strata onshore bornholm have been described in papers published over the last 150 years and constitute a classical topic in danish geology. however, previous studies have primarily focused on biostratigraphical aspects. contemporaneous strata throughout most of central and western denmark are deeply buried and hence, have only been reached by a few deep exploration wells (fig. 1), and thus the succession is rather poorly known. only in slagelse-1 and terne-1 in kattegat have the palaeozoic shales been fully penetrated whereas only the topmost parts of the silurian shales were reached by the rønde-1, nøvling-1, pernille-1 and stina-1 wells (fig. 1). a new shale gas exploration well, vendsyssel-1, is planned to be drilled in 2015 and is expected to penetrate the entire lower palaeozoic shale sequence in northern jylland (fig. 1). the sommerodde-1 well th e well was drilled as close as technically possible to the sommerodde beach locality described by bjerreskov & jørgensen (1983) where the youngest silurian strata onshore bornholm are exposed at low tide. prior to drilling, a seismic section was recorded in april 2012 and relevant water well sweden lower palaeozoic strata caledonian front borehole ringkøbing–fyn high pernille-1 kattegat terne-1 vendsyssel-1 rønde-1 nøvling-1 slagelse-1 stina-1 denmark norwegian–danish basin 50 km germany bornholm fig. 2 jylland fig. 2. geological map of southern bornholm showing the positions of scientific wells drilled 2005–2012. borggård-1 (dgu 247.627), skelbro-2 (dgu 246.817), billegrav-2 (dgu 248.61), hjulmagergård-1 (dgu 246.838) and sommerodde-1 (dgu 248.62) are shown as red dots, and the re-logged water wells sømarken-3 (dgu 248.39), sømarken-4 (dgu 247.312) and golfbanen-1 (dgu 248.54) are shown as blue dots. areas where 2d seismic data were acquired are shown as black rectangles. the geological map is from a compilation by graversen (2009). nexø sommerodde-1 billegrav-2 golfbanen sømarken-4 sømarken-3 borggård-1 precambrian mesozoic palaeozoic fault 2 km 15°5´e 54°59.2´n hjulmagergård-1 skelbro-2 fig. 1. distribution of lower palaeozoic strata and deep wells reaching at least top silurian in denmark north of the caledonian front. the vendsyssel-1 well is planned to be drilled in 2015. © 2015 geus. geological survey of denmark and greenland bulletin 33, 9–12. open access: www.geus.dk/publications/bull 1010 0 20 40 60 80 100 120 140 160 180 200 220 240 flow % fluid conductivity (ms/m) 0 40 80 120 160 3 4 510 100 1000 0 0 500 100 0 400 800 1200 r.m. a lu m s h al e d ic e llo gr . l in d eg . log unit l.c. m .c . u p p er o rd o vi ci an l o w er s ilu ri an fu ro n . q r as tr it es s h al e p-wave velocity (km/s)formation resistivity (ohm m)gamma ray (cps)optvlithologyperiod stratigraphy l.o. c yr to gr ap tu s sh al e casing (scale change) d1 d2 d3 a b2 b3 b4 g1 g5 g4 g2 e3 e1 e2 f1 f5 f4 f3 f2 g3 n.m. depth (m) grey shale limestonesandstoneblack shale dark grey shale light grey shale fig. 3. selected logs measured in the sommerodde-1 well. the arrows indicate significant water inflow zones. the optic televiewer (optv) picture of the well is a 360° view of the borehole wall. flow %: water flow scaled to the total flow of 1.4 m3/h. q: quaternary. lo: lower ordovician. furon.: furongian. mc: middle cambrian. lc: lower cambrian. lindeg.: lindegård formation. dicellogr.: dicellograptus shales. k: komstad limestone. rm: rispebjerg member. nm: norretorp member of the læså formation. cps: counts per second. 11 reports were examined to locate potential faults in the area (fig. 2). th e well was drilled in november 2012, well logging was done by both the geological survey of denmark and greenland and rambøll and the hole was plugged and permanently abandoned in november 2013. a total of ten log runs were made including a full wave sonic log and an optic televiewer recording that provided a high resolution image of the borehole wall (fig. 3). th e sommerodde-1 well cored silurian shales between 6.5 and 174.6 m, ordovician shale above the alum shale formation between 174.6 and 217.3 m, a thin komstad limestone (middle ordovician) between 217.3 and 217.9 m, the cambro-ordovician alum shale formation between 217.9 and 245.8 m and lower cambrian sandand siltstone between 245.8 and 250.3 m (fig. 3). a total of 168.1 m silurian shales were thus cored. th e lower silurian rastrites shale was completely cored and is 76.4 m thick in the well. th e alum shale formation is 27.9 m thick and includes the middle cambrian andrarum and exsulans limestone beds. th e well was terminated at 250.3 m in the norretorp member of the læså formation aft er having penetrated the rispebjerg member (4.5 m thick). log stratigraphy in the sommerodde-1 well pedersen & klitten (1990) established a detailed log stratigraphical scheme (labelled units a–g on figs 3, 4) based on the gamma-ray variation in the lower palaeozoic shales on bornholm. th e log stratigraphy permitted correlation of un-cored water wells with fully cored scientifi c wells. th e log stratigraphy was later emended to also include the resistivity log, from which more detailed lithological information can be gained (schovsbo et al. 2011). in the sommerodde-1 all log-based units defi ned by pedersen & klitten (1990) were identifi ed and the previously un-divided g unit was divided into four new log units, labelled g1–g4 (fig. 3). th e units a–f were described by schovsbo et al. (2011) and, hence, only the new log units are briefl y described here. th e g unit is correlated here with the cyrtograptus shale as was originally done by pedersen & klitten (1990). th e base of the g unit has not previously been cored nor recognised in logs from water wells on bornholm. it is here placed at 98.2 m; above this level the resistivity and sonic velocity log readings increase (figs 3, 4). th e g1 subunit is 24 m thick and consists of light green to dark grey shales. th e unit is characterised by a steady increase in resistivity readings. subunit g2 is 18.5 m thick and consists of light grey to dark green shale. subunit g3 is 16 m thick and consists of light grey to dark green shale. th e top of the g3 subunit is defi ned where an increase occurs in the resistivity log, refl ecting a change to darker lithologies. numerous bentonite beds characterised by low resistivity and high gamma ray readings occur in g2 and g3. subunit g4 is a 10 m thick, dark-coloured interval. th e subunit is readily identifi ed in the core and on the optic televiewer log and by a slightly lower gamma-ray log response and distinctly higher resistivity readings compared to the subunits above and below. th e top of the g4 subunit is defi ned at a point of decrease in the resistivity log. subunit g5 is 23.5 m thick and comprises light grey to green shale characterised by low and stable gamma-ray values. th e subunit includes numerous silty to sandy beds similar to those exposed on the beach just south of the well location (bjerreskov & jørgensen 1983). gamma ray (cps) 40 80 120 160 log units gamma ray (cps)gamma ray (cps) sommerodde-1 resistivity (ohm m) 100 1000 resistivity (ohm m) 10 100 1000 billegrav-2billegrav-2 d ep th (m ) 0 20 40 60 80 100 120 140 160 180 0 40 80 120 160 50 100 150 200 20 40 60 d ep th (m ) d ep th (m ) 0 20 40 60 golfbanen-1 g2 g1 g5 g4 g3 f2 f1 f3 f4 f5 e3 fig. 4. correlation of the silurian sections in the sommerodde-1, billegrav-2 and golfbanen-1 wells based on the gamma-ray (black line) and resistivity log (red line) responses. grey intervals show intervals where the log response was measured through the steel casing. 1212 correlation of silurian shales on southern bornholm th e g unit was originally defi ned based on the gammaray signature in the un-cored water well golfbanen-1 that penetrated a section of the cyrtograptus shale (pedersen & klitten 1990). as part of the present study the golfbanen-1 well was re-logged in order to expand the log signature and to enable correlation with the sommerodde-1 well (fig. 4). based on comparison with the sommerodde-1 well, the gamma ray and resistivity logs show that the golfbanen-1 well penetrated a sequence from the g3 to the uppermost f5 unit. th e sommerodde-1 well is also correlated with the fully cored billegrav-2 well that penetrated the succession from the lower part of the f5 unit and downwards, thereby closing the correlation gap between the billegrav-1 well and the exposures in the øle å water course as presented by pedersen & klitten (1990). several of the rastrites shale log units are slightly thicker in the sommerodde-1 well than in the billegrav-2 well (fig. 4). in the latter several fault zones occur in this interval and it appears that parts of the section are missing due to faulting. in the sommerodde-1 well no fault zones are identifi ed and it is believed that the recorded thicknesses refl ect the true stratigraphical thicknesses. water-flow properties th e water infl ow in sommerodde-1 was evaluated from an impeller fl ow log during discharge of 1.4 m3/h supplemented by the fl uid conductivity log measured during the same discharge from the well. together the two logs show that most of the infl ow took place within the uppermost 75 m of the well (fig. 3). th e infl ow rates decrease rapidly with depth and the section below 80 m contributes with less than 20% of the total infl ow into the well. th e water conductivity is highest in the alum shale formation and in the lower part of the rastrites shale, 350–400 ms/m, but from a depth of 125 m and upwards several small infl ow sources with lower conductivity stepwise reduce the conductivity of the upwards fl owing water until it has a conductivity of 50 ms/m (fig. 3). in spite of the many infl ow points the specifi c yield is quite low, less than 0.1 m3/h/m, suggesting that only a few of the fractures in the shales are open and that these fractures occur within the uppermost 75 m. similar conditions are seen in many wells penetrating danian limestone and cretaceous chalk found near the surface, where fracture-based hydraulic conductivity is mainly found in the upper part, less than 70 m below the prequaternary surface. th is is probably because the fractures were formed due to pressure release in connection with the last deglaciation. in contrast, the deep, water-bearing open fractures on bornholm, as for instance observed in the nexø formation in the borggård-1 well at a depth of c. 300 m, were probably formed due to tectonic uplift . conclusions new knowledge on the palaeozoic geology of bornholm has been gained since 2005 from fi ve fully cored and geophysically logged scientifi c wells, re-logging of some water wells and acquisition of seismic data. a complete section of the lower palaeozoic has been pieced together by correlation of overlapping scientifi c well sections. th e research on bornholm has provided detailed insight into the palaeozoic stratigraphy and established a lithoand log-stratigraphical frame that is applicable in a regional context. combined with ongoing drilling activities in northern jylland this will provide a much better understanding of the evolution of the palaeozoic in denmark. acknowledgements we are thankful to landowner jette staberg, peter turner of faxe kalk a/s, kurt nielsen of rambøll and klaus bauer of gfz. th e greater part of the funding was received from geocenter denmark and total e&p denmark b.v. references bjerreskov, m. & jørgensen, k.å. 1983: late wenlock graptolite-bearing tuff aceous sandstone from bornholm, denmark. bulletin of the geological society of denmark 31, 129–149. graversen, o. 2009: structural analysis of superposed fault systems of the bornholm horst block, tornquist zone, denmark. bulletin of the geological society of denmark 57, 25–49. lassen, a. & th ybo, h. 2012: neoproterozoic and palaeozoic evolution of sw scandinavia based on integrated seismic interpretation. precambrian research 204–205, 75–104. nielsen, a.t., klitten, k. & hansen, h.p.b. 2006: borggård-1: en ny stratigrafi sk kerneboring på bornholm. geologisk nyt 6, 4–10. pedersen, g.k. & klitten, k. 1990: anvendelse af gamma-logs ved korrelation af marine skifre i vandforsyningsboringer på bornholm. dansk geologisk forening årsskrift 1987–89, 21–35. schovsbo, n.h., nielsen, a.t., klitten, k., mathiesen, a. & rasmussen, p. 2011: shale gas investigations in denmark: lower palaeozoic shales on bornholm. geological survey of denmark and greenland bulletin 23, 9–13. authors’ addresses n.h.s. & k.k., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: nsc@geus.dk a.t.n., natural history museum of denmark. øster voldgade 5-7, dk-1350 copenhagen k, denmark. geological survey of denmark and greenland bulletin 41, 2018, 13-16 13 this study is based on a feasibility study for the danish energinet.dk to identify potential formations for brine storage near the gas storage facility at lille torup, northern jylland, denmark (fig. 1; hjuler et al. 2017). located on top of a salt structure, the gas storage facility comprises seven caverns, which have been washed out by circulating water in the salt dome. one cavern contains c. 520.000 m3 of intrusive brine that must be disposed of in order to increase the storage volume for gas. one option is to inject the brine into the subsurface if a target with appropriate storage properties can be identified, but it is a prerequisite that the stored brine does not compromise freshwater reservoirs. due to cost considerations, the brine storage should be situated within a radius of 50 km of the gas storage facility and at a depth not exceeding 2000 m. based on the national geothermal research conducted during the last decade, a number of sandy formations are considered potential storage reservoirs (fig. 2; e.g. mathiesen et al. 2009; vosgerau et al. 2016). around lille torup, these include the bunter sandstone/skagerrak, gassum, haldager sand and frederikshavn formations where the two former formations are discarded due to present-day burial depths exceeding 2000 m. in addition, the chalk group is considered a potential storage formation due to its importance as a hydrocarbon reservoir in the north sea, however, due to risk of leakage to the younger sediments and risk of environmental issues, the chalk was discarded as potential storage zone. geological background the lille torup area is located centrally in the danish basin, where the upper permian–mesozoic succession is 5–5.5 km thick. the basin was formed in the late carboniferous–early permian with basal rotliegendes coarse-grained clastic sediments and thick zechstein salts overlain by triassic sandstone, mudstone, carbonate rocks and salt (nielsen 2003). these are followed by lower jurassic mudstone, middle jurassic sandstone, upper jurassic–lower cretaceous mudstone and siltstone with few sandstone layers. the mesozoic succession terminates with c. 1200 m thick carbonate deposits. the salt structure at lille torup consists of mobilised zechstein salt penetrating the mesozoic succession. its top point is c. 250 m below the present-day surface. the haldager sand formation in the northern part of the basin is 2–150 m thick, but may exceed 200 m in rim synclines of salt structures, where sandstone commonly dominates the lithology. the frederikshavn formation is primarily present in the northern part of the basin and frequently includes sandstone layers. its thickness decreases southwards from 150–300 m to a few metres. the more than 1000 m thick chalkand limestone-dominated chalk group constitutes the topmost pre-quaternary formation in large parts of the danish basin. potential for brine storage near the gas storage facility at lille torup, northern jylland, denmark morten leth hjuler, morten sparre andersen, carsten møller nielsen, anders mathiesen, lars kristensen, nina skaarup and lars henrik nielsen r i n g k ø b i n g f y n h i g hn o r t h g e r m a n b a s i n n o r w e g i a n d a n i s h b a s i n 50 km lille torup gas storage facility salt diapir fault selected storage area 2d seismic line aarhus silkeborg randersviborg hobro skive aalborg holstebro 1 2 3 4 5 erslev-1 erslev-2 hyllebjerg-1 aars-1 farsø-1 kvols-1 hobro-1 gassum-1 skive-1 skive-2 mors-1 rødding-1 15 km fig. 1. the study area and selected storage areas within a radius of 50 km of the lille torup gas storage facility. © 2018 geus. geological survey of denmark and greenland bulletin 41, 13–16. open access: www.geus.dk/bulletin http://www.geus.dk/bulletin 1414 methods the subsurface within a radius of 50 km of the lille torup storage facility was screened for potential sandstone reservoirs suitable for storage. the local database comprises 11 vertical deep wells and an open grid of regional 2d seismic profiles of variable quality and resolution (fig. 1). some wells were excluded from the database due to location on top of salt structures (erslev-1–2 and skive-1–2), uncertain data quality (aars-1) or separation from the lille torup storage facility by fjord water (mors-1). the danish geothermal webgis application (vosgerau et al. 2016) provided maps of formation depth, formation thickness and potential reservoir sandstone thickness as well as reservoir parameters of relevant wells. the potential reservoir sandstone thickness map was developed for assessment of the geothermal potential and is used in this study as an indicator for injection capacity. reservoir properties derived from well logs include the depths of formation top and base, formation thickness, gross sand thickness (i.e. cumulated thickness of all sandstone layers), potential reservoir sandstone thickness (i.e. cumulated thickness of sandstone layers with a shale content <30% and a porosity >15%), as well as averaged values of porosity, permeability and transmissivity of the potential reservoir sandstone. for uncertainty considerations, see http://dybgeotermi.geus.dk/. potential formations for brine storage five potential storage areas were defined based on reservoir quality assessments obtained by integration of well-log data and webgis data improved with locally refined seismic interpretations. injectivity assessments were performed using eclipse 100 reservoir simulation software and petrel software. see hjuler et al. (2017) for details. the haldager sand formation (figs 3a–c) is presently buried more than 2000 m in large parts of the study area, but more shallow occurrences exist. the generally 50–150 m thick formation is dominated by sandstone known to be quartz-rich, which points to good reservoir properties. in areas of relatively shallow burial (<2000 m), the haldager sand formation may constitute a storage formation. the frederikshavn formation (figs 3d–f) is buried less than 2000 m and generally more than 100 m thick; it is thickest east of lille torup. several potential storage reservoirs with sufficient lateral extent can be identified. the frederikshavn formation constitutes a storage option. the chalk group (figs 3g, h) is buried at 100–700 m depth and is more than 1 km thick. on top of the lille torup salt structure, the salt movements may have fractured the c. 250 m of chalk, and increased permeability and thus reservoir quality. however, the overlaying quaternary deposits are not expected to possess sealing qualities and brine storage in the chalk could lead to environmental issues. the chalk group is therefore discarded as a potential storage formation. potential seals the chalk group outside the top of the lille torup salt structure is expected to effectively seal off pore water from the sandy formations beneath it due to its low permeability and great thickness. in addition, the clayey lower cretaceous unit is assumed to be of sufficiently low permeability to prevent pressure and pore-water propagation from below. reservoir parameters of the formations the haldager sand formation mainly comprises sandstone layers with porosities in the 18–22% range and permeabilities in the 140–360 md range (table 1). disregarding burial depth, the haldager sand formation is assumed to provide suitable storage properties. the sandstone layers of the frederikshavn formation have porosities in the 17–30% range, permeabilities in the 110–1500 md range and the thickness of potential reservoir sandstone in the 6–66 m range (table 1). in the kvols-1 well, however, the formation seems to have little or no storage potential, assumedly because clay minerals reduce both pore quaternary system lithostratigraphic unit post chalk group chalk group lower cretaceous unit reservoir reservoir reservoir reservoir assumed seal assumed seal assumed seal frederikshavn fm børglum fm flyvbjerg fm haldager sand fm fjerritslev fm gassum fm vinding fm oddesund fm tønder fm falster fm ørslev fm bunter sandstone fm zechstein group cretaceous jurassic triassic permian fig. 2. lithostratigraphic chart showing potential reservoirs (yellow) and assumed seals (brown). 15 space and permeability. the thickness of potential reservoir sandstone decreases from east to west (fig. 3f). suggested areas for brine storage the webgis-based maps (fig. 3) are suitable for assessing reservoir quality trends on a regional scale of tens of kilometres, but not on a local scale of up to a few kilometres as uncertainties will be significant. thus, the areas suggested for brine storage cover several square kilometres (fig. 1). area 1 includes the hyllebjerg-1 and farsø-1 wells (figs 1, 3), in which the haldager sand formation is evaluated to provide the better storage reservoir, with higher porosities and permeabilities than the frederikshavn formation (table 1). the two formations offer two storage options within a narrow depth interval. area 2 is situated above the rim syncline next to the salt structure beneath lille torup (figs 1, 3), where the frederikshavn and haldager sand formations may be up to 300–400 m thick (figs 3a, d) and include potential reservoir sandstone units more than 15 m thick at depths shallower than 1400 m. however, existing seismic data are insufficient to confirm these thickness estimates and the formation depths may be closer to 2000 m due to deposition in the rim synm below mean sea level thickness thickness h ald ag er s an d fm fr ed er ik sh av n fm c ha lk g ro up hy-1 a-1f-1 k-1 ho-1 g-1 m-1 hy-1 a-1f-1 k-1 ho-1 g-1 m-1 hy-1 a-1f-1 k-1 ho-1 g-1 m-1 hy-1 a-1f-1 k-1 ho-1 g-1 m-1 hy-1 a-1f-1 k-1 ho-1 g-1 m-1 hy-1 a-1f-1 k-1 ho-1 g-1 m-1 hy-1 a-1f-1 k-1 ho-1 g-1 m-1 hy-1 a-1f-1 k-1 ho-1 g-1 m-1 hy-1 a-1f-1 k-1 ho-1 g-1 m-1 hy-1 a-1f-1 k-1 ho-1 g-1 m-1 hy-1 a-1f-1 k-1 ho-1 g-1 m-1 hy-1 a-1f-1 k-1 ho-1 g-1 m-1 hy-1 a-1f-1 k-1 ho-1 g-1 m-1 hy-1 a-1f-1 k-1 ho-1 g-1 m-1 hy-1 a-1f-1 k-1 ho-1 g-1 m-1 hy-1 a-1f-1 k-1 ho-1 g-1 m-1 r-1 s-1 s-2r-1 s-1 s-2 r-1 s-1 s-2r-1 s-1 s-2 r-1 s-1 s-2r-1 s-1 s-2 r-1 s-1 s-2r-1 s-1 s-2 r-1 s-1 s-2r-1 s-1 s-2 above sea level ≥ 15 m 0–100 100–200 200–300 300–400 400–500 500–600 600–700 700–800 1200–1300 1300–1400 1400–1500 1500–1600 800–900 900–1000 1000–1100 1100–1200 not present 0–50 m 50–100 100–150 150–200 200–250 250–300 300–350 350–400 400–450 450–500 500–550 550–600 600–650 > 700 650–700 not present at 800–3000 m below mean sea level 50 km lille torup gas storage facility salt diapir fault wells: study area e-1 erslev-1 e-2: erslev-2 f-1: farsø-1 g-1: gassum-1 ho-1: hobro-1 hy-1: hyldebjerg-1 k-1: kvols-1 m-1: mors-1 r-1: rødding-1 s-1: skive-1 s-2: skive-2 a-1: aars-1 depth to formation formation thickness potential reservoir sand thickness a b c d e f g h fig. 3. webgis-generated maps of potential brine storage formations in the study area. reservoir quality is assessed based on depth to formation top, formation thickness and potential reservoir sand thickness. 1616 cline. thus, the estimated reservoir parameters are based on average values of the nearby wells, hyllebjerg-1, farsø-1, hobro-1 and kvols-1. area 3 includes the hobro-1 well (figs 1, 3), in which the haldager sand and frederikshavn formations are estimated to offer two storage options within a narrow depth interval, the haldager sand formation providing the better reservoir (table 1). area 4 includes the gassum-1 well (figs 1, 3). the frederikshavn formation is more shallowly buried in this area (<1000 m; fig. 3d) and offers the most excellent storage properties in the study area (table 1). the haldager sand formation is not present in area 4 (fig. 3a). area 5 includes the kvols-1 well (figs 1, 3). this well indicates good reservoir properties of the haldager sand formation at a burial depth of 1940–1955 m, but also that the frederikshavn formation is a poor reservoir (table 1). assessment of injection rate and pore pressure indicative injection rates (shown as a well injection index) and pore pressures for the frederikshavn and haldager sand formations in the entire study area are presented in table 1. reservoir parameters are calculated as averages of all wells. the well injection index is assumed to correspond to the well production index; the pressure of the subsurface pore fluid in the formations is assumed to be hydrostatic. conclusions the frederikshavn and haldager sand formations constitute potential brine storage formations in the larger lille torup area and are sealed off by the chalk group and lower cretaceous unit. the haldager sand formation offers the best reservoir properties but at greater depth than the frederikshavn formation. the well injectivity index and pore pressure indicate favourable conditions for brine storage. five potential storage areas are suggested, compromising between burial depth and distance from the gas storage facility. references hjuler, m.l., andersen, m.s., nielsen, c.m., mathiesen, a., kristensen, l., skaarup, n. & nielsen, l.h. 2017: potential for establishing an injection well for brine storage near the gas storage facility at lille torup. danmarks og grønlands geologiske undersøgelse rapport 2017/20, 34 pp. mathiesen, a., kristensen, l., bidstrup, t. & nielsen, l.h. 2009: vurdering af det geotermiske potentiale i danmark. danmarks og grønlands geologiske undersøgelse rapport 2009/59, 85 pp. nielsen, l.h. 2003: late triassic – jurassic development of the danish basin and the fennoscandian border zone, southern scandinavia. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 459–526. vosgerau h., mathiesen a., andersen m.s., boldreel l.o., hjuler m.l., kamla e., kristensen l., pedersen c.b., pjetursson b. & nielsen l.h. 2016: a webgis portal for exploration of deep geothermal energy based on geological and geophysical data. geological survey of denmark and greenland bulletin 35, 23–26. table 1. reservoir parameters of the haldager sand and frederikshavn formations in five areas suggested for brine storage reservoir parameters based on the vertical wells in the suggested areas (fig. 1). reservoir parameters of area 2 based on average values of the hyllebjerg-1, farsø-1, hobro-1 and kvols-1 wells. for detailed descriptions of parameters and uncertainties, see http://dybgeotermi.geus.dk/. area 1 farsø-1 haldager sand 1934 1952 18 16 11 18 138 2 hyllebjerg-1 haldager sand 1885 1894 9 9 9 22 344 3 average haldager sand 1909 1923 14 12 10 20 241 3 farsø-1 frederikshavn 1689 1839 150 25 9 17 106 1 hyllebjerg-1 frederikshavn 1664 1810 146 37 15 19 188 3 average frederikshavn 1676 1824 148 31 12 18 147 2 area 2 average haldager sand 1900 1920 20 16 12 20 255 3 average frederikshavn 1700 1756 56 33 8 15 133 1 area 3 hobro-1 haldager sand 1852 1891 39 24 18 19 175 3 hobro-1 frederikshavn 1741 1806 65 12 6 19 238 1 area 4 gassum-1 haldager sand 1176 1178 2 2 2 19 188 0.5 gassum-1 frederikshavn 1053 1154 101 66 65 30 1500 97 area 5 kvols-1 haldager sand 1940 1955 15 15 9 21 363 3 kvols-1 frederikshavn 1856 1912 56 56 0 5 0 0 well average haldager sand 1757 1774 17 13 10 20 242 2 9 well average frederikshavn 1601 1704 104 39 19 28 406 8 27 well formation thickness property to p be lo w m ea n se a le ve l ( m ) ba se b el ow m ea n se a le ve l ( m ) fo rm at io n th ick ne ss (m ) g ro ss sa nd (m ) po te nt ial re se rv oi r s an ds to ne (m ) av er ag e po ro sit y (m d ) av er ag e re se rv oi r p er m ea bi lit y (m d ) re se rv oi r t ra ns m iss ivi ty (m d ) w el l i nj ec tio n in de x (m 3 /b ar /d ay ) authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen, denmark. e-mail: mlh@geus.dk. mailto:mlh@geus.dk geological survey of denmark and greenland bulletin 33, 2015, 21-24 21 consistency of postglacial geodynamics for the kattegat region, southern scandinavia, based on seismological, geological and geodetic data søren gregersen and peter h. voss th e earthquake map of denmark is constantly being improved. together with data from western sweden and southern norway it shows more and more convincingly a gradual, scattered earthquake activity across the kattegat region from low activity in the precambrian basement of scandinavia to lack of earthquakes in south-western denmark and northern germany. th e activity is only partly connected with mapped geological features. th e three most recently felt earthquakes in denmark augment and support this pattern with two or three activity concentrations in the seas around denmark (fig. 1). th e smoothness and irregularities of this picture must in some way be related to the geological structure as well as to the geodynamic pattern of postglacial uplift mapped from geology and geodesy. since the dominant stress fi eld, from the lithospheric plate motion is smooth (gregersen & voss 2010), a natural question is whether the picture © 2015 geus. geological survey of denmark and greenland bulletin 33, 21–24. open access: www.geus.dk/publications/bull 8°e 12°e 16°e 20°e 56°n zone ringkøbing – fyn high tornquist sweden germany denmark norway kattegat magnitude 100 km important geological line noticeable historical earthquake latest felt earthquakes 5 4 3 2 magnitude fig. 1. the latest three felt earthquakes in denmark shown on the best available earthquake map of the kattegat and the surrounding regions (gregersen & voss 2014). little earthquake activity is seen north-east of the gradual margin of the fennoscandian shield between the tornquist zone and the ringkøbing–fyn high. in contrast, no activity is seen in south-western denmark and north-western germany. only a few small earthquakes have been recorded in the south-eastern part of the area shown on the map; this is because only few seismographs are located in that region. 2222 of postglacial uplift , seen from the geological and geodetic perspective shows local irregularities? th e broad-scale uplift pattern in denmark since the last ice age has been known since the review published by mertz (1924). all later papers have referred to this general pattern of uplift towards the north-east and subsidence towards the south-west with reference to the present sea level. th e present paper relates the regional seismological pattern to the smooth and consistent patterns in postglacial geology and geodesy, as well as to the existing local investigations on geodynamics. interpretations of the measured uplift of shorelines on læsø, in the middle of the kattegat (fig. 2), have given rise to opposing views on regularity versus irregularity between the present authors and hansen et al. (2012). th e measurements show that the elevation of the oldest postglacial shorelines on læsø fi t in the regional pattern with regard to elevation, but not with regard to time. in this paper we propose that the time diff erence is only apparent. uplift patterns recent overviews of the geological evolution of the region of denmark, including the uplift since the time of fl ooding, termed the littorina sea transgression, have been presented by christensen (2001) and noe-nygaard et al. (2006). th e former provided ages for the oldest shorelines, from radiocarbon dating, pollen and human settlement evidence. it is seen that the uplift increases gradually from the south-west towards the north-east (fig. 2). th e general pattern with increasing ages towards the north-east has been convincingly argued for by christensen (2001). th e ages in fig. 2 are the oldest for each location for the littorina sea period that lasted several thousands of years. during this long time period the postglacial uplift of the lithosphere approximately followed the rising water level, within about one metre. both were infl uenced by the melting ice sheets aft er the last ice age. th e situation, with repeated small transgressions and regressions, was illustrated by christensen (2001), noe-nygaard et al. (2006) and clemmensen et al. (2012). th e latter paper shows that this period of several thousand years ends rather abruptly with a large regression lasting until present time. th e start of the regression can be dated by optically stimulated luminescence (osl), when no later transgressions disturb the dark, covered parts of the beach sediments used for dating. th e elevation of raised beaches and ages determined for anholt (fig.  2) by clemmensen et al. (2012) are interpreted as a period of ‘simultaneous’ land uplift and sea-level rise for several thousands of years followed by a period of continuing land uplift and slower sea-level changes. th is is suggested generalised for the kattegat region in fig. 3 with times that are diff erent for diff erent locations. in fig. 3 two kinds of arrows are seen. th e thick arrow points to the oldest shoreline that has not been disturbed by a transgression. sediments from this shoreline can be dated by the osl method. on anholt it has been dated to 2300 years bc (clemmensen et al. 2012). th e thin arrows point to archaeological time periods which can be dated by pollen hobrohobro 7.57.5 4500 bc4500 bc aalborgaalborg hobrohobrohobrohobro aalborg denmarkdenmark jyllandjylland denmark 50 km10°e 10°e 56° 56°n hobro copenhagen 23 6300 bc 20 6000 bc 15 12 5200 bc 9 4600 bc 8 7.5 4500 bc 4.8 4200 bc 1 3600 bc 0 3600/2500 bc 0 2 2 6 8 10 12 4 4 2 3 frederikshavn læsø anholt rügen jylland skåne skagerrak kattegat aalborg sweden fig. 2. general picture of the elevation patterns of postglacial shorelines in metres above present sea level, according to christensen (2001). shaded circles and zones refer to locations mentioned in the text, where irregularities have been claimed. ages of highest shorelines of the littorina sea, i.e. after the last ice age, are from christensen (2001). this figure is a corrected version of a figure presented by gregersen & voss (2012) and by gregersen (2014), where the ages were incorrectly quoted. 23 analysis, radiocarbon dating or human settlement indicators for places close to the beach where plants or human remains have survived close to the shoreline or just above it. in fig. 2 the anholt age from christensen (2001) is 4500 years bc, i.e. 2200 years before the osl age mentioned above. only in special cases will an older shoreline from within the littorina sea time, protected from transgression, show up by osl dating as in anholt (clemmensen et al. 2012). on læsø all ages were determined by osl and all are young compared to regions with similar elevations of the highest shorelines (fig. 2; hansen et al. 2012; j.m. hansen, personal communication, 2013–2014). th ose authors also suggest that older shorelines are not observed on læsø because they were eroded by waves. we suggest that this shows that only the right part of fig. 3 is relevant for læsø. in contrast, data from both the left and right parts are available from anholt. many other regions such as northern jylland and western sweden are well described by the left part of fig. 3 (christensen 2001, fi g. 2; christensen & nielsen 2008). th is point was discussed by gregersen & voss (2010, 2012, 2014) and gregersen (2014) who suggested that the age determinations ought to be redone using the same methods for points of comparable uplift in northern jylland, on læsø and in western sweden (see fig. 2). a smooth picture with only small irregularities is also shown by a recent geodetic picture of uplift velocities based on all available data from levelling, global positioning system (gps) and sea-level gauges (fig. 4). both the geological picture illustrated by fig. 2 and the geodetic picture in fig. 4 show the main, general features. geological and physically small irregularities are generalised into the broad-scale regional picture. in addition, the geodetic picture is based on measurements carried out over only approximately 100 years. updating with greater detail will be appreciated in the future. figures 2 and 4 indicate regularity so far. discussion of irregularities from the literature in addition to the irregularity proposed for læsø, several more irregularities have been proposed for diff erent regions in denmark (figs 2, 4). some of them show irregularities in the sedimentary succession which may be connected to larger-scale postglacial faulting, e.g. the carlsberg fault in the copenhagen area described by rosenkrantz (1937) and ovesen et al. (2002) or faulting reported from mid-jylland by jakobsen & pedersen (2009). other irregularities could be directly related to the postglacial uplift pattern, as described by lykke-andersen & borre (2000) and gregersen & schmidt (2001) for the region from hobro via aalborg to frederikshavn or by hansen et al. (2012) for harbours in time of littorina sea h e ig h t o f b e a ch p re se n t ti m e c. 1 m e tr e increasing age 2.0 1.8 1.6 1.4 1.2 1.2 1.0 0.8 0.6 0.6 0.4 2.2 25 km denmark sweden læsø 0.8 southern denmark hallandsåsen copenhagen fig. 3. conceptual model showing the elevation of shorelines in kattegat as a function of age. the model shows the generalised concept of uplift with minor oscillations in the grey band of the order of one metre from several sources mentioned in the text. the arrows point to age data, thin arrows to archaeological data, thick arrow to osl data. only for anholt older osl data supplement the archaeological dates. fig. 4. generalised uplift velocities (mm/year) from knudsen et al. (2012) based on levelling, gps and sea-level gauges. the velocities have been corrected from an earlier version by gregersen & voss (2014). the shaded circles and zones show locations where irregularities have been claimed, and which are discussed in the text. 2424 southern denmark, based on 100 years of water-level measurements. th e irregularity locations in the middle of skåne in southern sweden and on rügen in northern germany are based on geodetic measurements. th e case in southern sweden (pan et al. 1999) has been rejected by swedish geodesists because of inaccuracy (h.-g. scherneck, personal communication 2015). th e case in northern germany is based on a personal communication by r. dietrich in 2008. it was claimed that geodetic levelling indicates similar changes in uplift rates as those seen in northern jylland by lykke-andersen & borre (2000), but it awaits further confi rmation by geodesists. th e claims by mörner (2003) on postglacial earthquakes near hallandsåsen, based on geological irregularities were taken up on a fi eld excursion in the summer of 2013 by 12 international seismologists on the initiative of the fi rst author of this article. th e outcome was “possible but not probable”. conclusion th e discussion as to whether the evidence for lack of major irregularities in the geology and geodesy fi elds is suffi ciently strong to be signifi cant is heated (hansen et al. 2012; gregersen & voss 2012). recent papers by gregersen & voss (2012, 2014) and by gregersen (2014) claim that the irregularities are so small compared to the uncertainties in the measurements that they are insignifi cant, and thereby not disturbances of the regularity. th ey are only possibilities. th ey need supporting evidence in the form of irregularities in the generalised geodynamics. our view is that such supporting evidence is lacking. th e consequence of the consistency of the regional geophysical fi elds is that the short term, thousand-years evidence from seismology adequately describes the geodynamics and its irregularities. we suggest that the use of diff erent methods for dating of the shorelines results in that diff erent phases of the uplift are dated. th e osl dates refl ect the last infl uence of the waves on the loose sediments of the shore, whereas the archaeological dates refl ect the several thousand-year long littorina sea time when land plants left remains in the dry ground near the shoreline. th e age determinations are diff erent because diff erent phenomena were measured. references christensen, c. 2001: kystbosættelse og havniveauændringer i stenalderen. in: jensen, o.l., sørensen, s.a. & hansen, k.m. (eds): danmarks jægerstenalder – status og perspektiver, 183–193. hørsholm: hørsholm egns museum. christensen, c. & nielsen, a.b. 2008: dating the littorina sea shore levels in denmark on the basis of data from a mesolithic coastal settlement on skagens odde, northern jutland. polish geological institute special papers 23, 27–38. clemmensen, l., nielsen, l. & konnerup-madsen, j. 2012: anholt. geoviden 2012/1, 20 pp. gregersen, s. 2014: jordskælvsrisiko i danmark? – forslag til fremtidens studier. kvant 2014/2, 30–34. gregersen, s. & schmidt, k. 2001: tektonik i danmark. sorgenfrei–tornquist zonen. geologisk nyt 2001/1, 16–17. gregersen, s. & voss, p. 2010: irregularities in scandinavian postglacial uplift /subsidence in time scales tens, hundreds, thousands of years. journal of geodynamics 50, 27–31. gregersen, s. & voss, p.h. 2012: eff orts to include geological and geodetic observations in the assessment of earthquake activity in denmark. geological survey of denmark and greenland bulletin 26, 41–44. gregersen, s. & voss, p.h. 2014: review of some signifi cant claimed irregularities in scandinavian postglacial uplift on timescales of tens to thousands of years – earthquakes in denmark? solid earth 5, 109–119. hansen , j.m., aagaard, t. & binderup, m. 2012: absolute sea levels and isostatic changes of the eastern north sea to central baltic region during the last 900 years. boreas 41, 180–208. jakobsen, p.r. & pedersen, s.a.s. 2009: fracture valleys in central jylland – a neotectonic feature. geological survey of denmark and greenland bulletin 17, 33–36. knudsen, p., engsager, k. & khan, s.a. 2012: dokumentation for beregning af ny uplift -model 2011, 19 pp. unpublished report. lyngby: dtu space. lykke-andersen, h. & borre, k. 2000: aktiv tektonik i danmark – der er liv i sorgenfrei–tornquist zonen. geologisk nyt 2000/6, 12–13. mertz, e.l. 1924: oversigt over de senog postglaciale niveauforandringer i danmark. danmarks geologiske undersøgelse ii. række 41, 49 pp. mörner, n.-a. 2003: paleoseismicity of sweden – a novel paradigm, 320 pp. stockholm: university of stockholm. noe-nygaard, n., knudsen, k.l., & houmark-nielsen, m. 2006: fra istid til og med jægerstenalder. in: sand-jensen, k. & larsen, g. (eds): naturen i danmark, geologien, 303–331. copenhagen: gyldendal. ovesen, n.k., blem, h., gregersen, s., møller, h.m.f. & frederiksen, j.k. 2002: recente terrænbevægelser i københavn. dansk geoteknisk forenings bulletin 19, 183–192. pan, m., sjöberg, l.e., talbot, c. & asenjo, e. 1999: gps measurements of crustal deformation in skåne, sweden between 1989 and 1996. gff 121, 67–72. rosenkrantz, a. 1937: bemærkninger om det østsjællandske daniens stratigrafi og tektonik. meddelelser fra dansk geologisk forening 9, 199–212. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: sg@geus.dk geological survey of denmark and greenland bulletin 33, 2015, 41-44 41 the continental shelf project of the kingdom of denmark – status and issues christian marcussen, finn mørk, thomas funck, willy lehmann weng and mikael pedersen th is paper summarises the status of the continental shelf project of the kingdom of denmark aft er the recent submission for an extended continental shelf in the area to the north of greenland. we discuss some of the similarities between the submission areas north of the faroe islands and north of greenland including the morphological continuation of ridges extending seaward of the geomorphical continental shelf. documentation of the sediment thickness in the adjoining basins and sediment continuity with the continental slope plays a vital role in the delineation of the outer limits of the extended continental shelf. here, we compare how these issues were addressed around the well-studied faroe islands and in the sparsely surveyed arctic ocean. th e kingdom of denmark ratifi ed the 1982 united nations convention on the law of the sea (unclos) in 2004. according to annex ii of unclos, the particulars of the outer limits of the extended continental shelf beyond 200 nautical miles shall be submitted to the commission on the limits of the continental shelf (clcs) within a period of ten years aft er ratifi cation. to acquire the necessary data for delineating the extended continental shelf, the continental shelf project of the kingdom of denmark was launched by the ministry of higher education and science in cooperation with the government of the faroe islands and the government of greenland (marcussen et al. 2004; marcussen & heinesen 2009). five partial submissions have been submitted by the kingdom of denmark regarding areas north of the faroe islands in 2009, south of the faroe islands in 2010, south of greenland in 2012, north-east of greenland in 2013 and north of greenland in 2014 (fig. 1). the area north of the faroe islands north of the faroe islands, the extinct ægir sea-fl oor spreading ridge is a prominent feature of the continental margin (figs 1, 2). th e initial opening of the north atlantic occurred along the ægir ridge with active sea-fl oor spreading in the eocene and oligocene. subsequent thermal subsidence of the oceanic crust resulted in the formation of the northern deep with a sediment accumulation of up to 3 km. th e clcs examined the partial submission regarding the area north of the faroe islands between 2012 and 2014 and gave special attention to: (1) the morphological continuation of the ægir ridge and (2) the continuity of sediments throughout the northern deep. in a strict morphological sense, the land mass of the faroe islands is connected to the faroe–iceland ridge (fig. 2). furthermore, the faroe–iceland ridge coalesces with and is morphologically linked to the ægir ridge. th is means that the ægir ridge is morphologically continuous with the rest of the continental margin © 2015 geus. geological survey of denmark and greenland bulletin 33, 41–44. open access: www.geus.dk/publications/bull nfm 150°e 1 000 km 30°w 50°n 30°e 50°n 70 °n negm nfm sfm sgm sgm sgm ngm ru s s i a n fe d e r a t i o n no rway no rwayc a n a d a g r e e n l a n d u sa fa ro e is la nd s fig. 1. the five partial submissions of the kingdom of denmark. red line: the 200 nautical mile limit of the five arctic coastal states in the arctic ocean. negm: north-east greenland margin. nfm: north faroes margin, ngm: north greenland margin. sfm: south faroes margin. sgm: south greenland margin. colour scale: see fig. 3. 4242 and lies within a common envelope of the foot of the continental slope. th us, in the sense of unclos, the ægir ridge is an integral part of the continental margin of the faroe islands, notwithstanding the tectonic and crustal differences between the two terrains (recommendations of the commission 2014). th e sediment thickness in the northern deep was documented primarily using modern seismic refl ection data. selected seismic lines were re-processed in order to improve the defi nition of the base of the sediments as well as the determination of the seismic velocities used for the depth conversion. th e clcs agreed with the procedure used to establish the sediment thickness, i.e. the methodology of depth conversion and the seismic interpretations. however, the clcs initially disagreed with the method by which the kingdom of denmark demonstrated sedimentary continuity throughout the northern deep. to satisfy the clcs, additional seismic lines, a sediment thickness map and gravity maps were provided and subsequently deemed suffi cient for this purpose. in march 2014, the clcs adopted the recommendations regarding the partial submission north of the faroe islands. th e commission agreed with the determination of the fi xed points establishing the outer limits of the continental shelf north of the faroe islands as originally listed in the submission. the area north of greenland th e lomonosov ridge is a sliver of continental crust that extends for a distance of almost 1800 km across the arctic ocean (fig. 3). th e ridge is 45 to 200 km wide, mostly fl at-topped to slightly rounded at its crest and rises from water depths of more than 4300 m in the adjacent basins to typically 1000 to 1300 m. th e shallowest part of the ridge is found towards greenland. th e lomonosov ridge was separated from the barents and kara shelves during the paleocene (55 to 60 ma), when sea-fl oor spreading started to open the eurasia basin. th is process continues today along the ultraslow-spreading gakkel ridge. th e geological development of the amerasia basin is not well understood, but most authors agree that the alpha and mendeleev ridges in that basin comprise a large igneous province of cretaceous age (e.g. jackson et al. 1986) and that sea-fl oor spreading has occurred in the canada basin (e.g. grantz et al. 1998). th e lomonosov ridge is morphologically continuous with the lincoln shelf north of greenland and furthermore shares geological characteristics with the land mass of greenland. th ese geological characteristics are based on evidence for continuous continental crust (jackson et al. 2010), similar lithologies and ages of rocks sampled from the fl anks of the ridge compared to rocks of the adjacent land masses, and a common tectonic history. in the sense of unclos, this means that the lomonosov ridge is a submarine elevation that is a natural component of the continental margin of greenland and therefore the extended continental shelf north of greenland extends across the arctic ocean to the 200 nautical mile line from russia (executive summary 2014). facing the amerasia basin, the lomonosov ridge is morphologically amalgamated with the alpha and mendeleev ridges and the chukchi borderland, which together are regarded as one complex but coherent morphological sea-fl oor high relative to the adjacent deep ocean fl oor within the eurasia and amerasia basins. th at is, all of this complex sea-fl oor high lies within a common envelope of the foot of the continental slope. however, the existing geological data for some of the individual sea-fl oor highs are insuffi cient to prove that they share geological characteristics with the greenland land mass. for this reason, and in accordance with unclos, the extended continental shelf, based on the alpha and mendeleev ridges and the chukchi borderland, is constrained to a distance of 350 nautical miles from the territorial sea baseline of greenland (executive summary 2014). 4°w 68°n 64°n 62°n 8°w8°w 0°8°w 100 kmf a ro e s h e t l a n d c h a n n e l j a n m a y e n r i d g e f a r o e i s l a n d s f aro e ice land ridge v ø r i n g p l a t e a u æ g i r r i d g e nor thern d eep v ø r i n g p l a t e a u fig. 2. the northern continental shelf of the faroe islands showing the submission area. the continental shelf beyond 200 nautical miles in the area north of the faroe islands, as delineated in the submission, amounts to 87 792 km2 in area and is highlighted. green: agreed maritime boundaries within 200 nautical miles. red: faroese 200 nautical mile limit. yellow: iceland’s and norway’s 200 nautical mile limits. orange: outer limit of the continental shelf north of the faroe islands. colour scale: see fig. 3. 43 by analogy with the ægir ridge north of the faroe islands, bathymetric and morphological data show that the gakkel ridge coalesces with the continental slope north of greenland. th e oceanic, sea-fl oor-spreading characteristics of the gakkel ridge diff er geologically from the continental crust of greenland; hence the extended continental shelf based on the gakkel ridge cannot exceed 350 nautical miles from the territorial sea baseline of greenland (executive summary 2014). published seismic data from the amundsen basin north of greenland (weigelt & jokat 2001) indicate the presence of sediments of suffi cient thickness (i.e. at least 1200 m) for use in extending the continental shelf. for this reason, a purpose-built seismic acquisition system was developed (hopper et al. 2012) that could operate in the arctic sea ice and at sub-zero temperatures. to locate the most promising regions for seismic data acquisition, a sediment thickness map was derived from gravity inversion (døssing et al. 2014). th is study was based on the lomgrav airborne gravity data (døssing et al. 2013). th ree geus lomrog expeditions were carried out in 2007, 2009 and 2012 to acquire seismic data in the arctic ocean. however, ice conditions only allowed for a maximum streamer length of 300 m, which is too short to extract velocities from the seismic refl ection data. for this reason, sonobuoys were deployed along the profi les to record the seismic energy at larger off sets of up to 34 km. ray-tracing modelling was then used to obtain velocity models of sediments in the amundsen basin. th e velocities show little scatter (fig. 4). assuming normal compaction, the data points can be approximated by an exponential function of instantaneous slowness versus burial depth (al-chalabi 1997) using reduced major axis regression techniques. th e velocity–depth relationship obtained from this approach is shown in fig. 4 and can be used to convert the sediment thickness from time to depth. th e seismic refl ection lines surveyed were generally less than 50 km long and focused on regions that were more than 60 nautical miles away from the lomonosov ridge. as the clcs requires the documentation of sediment continuity, supplementary data had to be used. in ice-covered areas gravity data can supplement a sparse seismic database. for this reason, the results of the gravity inversion (døssing et x x x x x x x x x x x x x x x x x x x x x x x x x xx x x x x lena trough l a p t e v s h e l f m endeleev r idge a m e r a s i a b a s i n e u r a s i a b a s i n canada basin podvodnikov bas in kara shelf e a s t s i b e r i a n s h e l f g ak k e l r id ge amundsen basin barents shelf a l p h a r i d g e nansen basin r i d g e morris jesup rise makarov basin lincoln shelf l o m o n o s o v canadian arct ic archipelago shel f b e a u f o r t s h e l f c h u k c h i s h e l f c hukchi borderland greenland severnaya zemlya f ranz j osef land alaska svalbard ellesmere island siber ia north pole 1 6 0 °w 8 0 °e 2 0 °e 1 0 0 °e 400 km 0 m 1 000 m 2 000 m 3 000 m 4 000 m bathymetry 8 0 °n 8 0 °n 140°e180° 100°w 0°80°w 20°w60°w 75°n70°n fig. 3. bathymetric map showing the main features of the arctic ocean. 4444 al. 2014) were used to document continuity of the sediments supported by additional subbottom profi ler and seismic refl ection data. concluding remarks th e continental shelf project of the kingdom of denmark mapped fi ve areas relevant to extended continental shelf beyond 200 nautical miles. th e project gathered bathymetric, geodetic, geological and geophysical data to compile the submission documents required by the clcs. despite diffi cult logistics, challenging climatic conditions and permanent ice cover in some of the fi ve areas, the project was completed within the initially allocated budget of 330 million dkk (44 million €) and the given time frame. large data sets have been acquired in the fi ve submission areas and key results have been or will be published in scientifi c journals. particularly valuable are the datasets in the arctic ocean as this remains a poorly studied frontier region. furthermore, the project was able to support other science projects by sharing logistics, in particular by providing access to icebreakers. th e results could not have been achieved without extensive international cooperation. due to the modus operandi of the clcs, recommendations for the latest submission north of greenland cannot be expected before the middle of the next decade. it represents a signifi cant challenge to keep both data and knowledge within the project updated and state of the art. references al-chalabi, m. 1997: instantaneous slowness versus depth functions. geophysics 62, 270–273. døssing, a., hopper, j.r., olesen, a.v., rasmussen, t.m. & halpenny, j. 2013: new aero-gravity results from the arctic: linking the latest cretaceous–early cenozoic plate kinematics of the north atlantic and arctic ocean. geochemistry, geophysics, geosystems 14, 4044–4065. døssing, a., hansen, t.m., olesen, a.v., hopper, j.r. & funck, t. 2014: gravity inversion predicts the nature of the amundsen basin and its continental borderlands near greenland. earth and planetary science letters 408, 132–145. executive summary 2014: partial submission of the government of the kingdom of denmark together with the government of greenland to the commission on the limits of the continental shelf – th e northern continental shelf of greenland. http://www.un.org/depts/los/clcs_ new/submissions_ fi les/dnk76_14/dnk2014_es.pdf gernigon, l., gaina, c., olesen, o., ball, p.j., péron-pinvidic, g. & yamasaki, t. 2012: th e norwegian basin revisited: from continental breakup to spreading ridge extinction. marine and petroleum geology 35, 1–19. grantz, a. et al. 1998: phanerozoic stratigraphy of northwind ridge, magnetic anomalies in the canada basin, and the geometry and timing of rift ing in the amerasia basin, arctic ocean. geological society of america bulletin 110, 801–820. hopper, j.r., trinhammer, p., marcussen, c. & funck, t. 2012: acquisition of seismic data in ice fi lled waters. american geophysical union, fall meeting 2012, abstract #c13e-0680. jackson, h.r., forsyth, d.a. & johnson, g.l. 1986: oceanic affi nities of the alpha ridge, arctic ocean. marine geology 73, 237–261. jackson, h.r., dahl-jensen, t. & the lorita working group 2010: sedimentary and crustal structure from the ellesmere island and greenland continental shelves onto the lomonosov ridge, arctic ocean. geophysical journal international 182, 11–35. marcussen, c. & heinesen, m. 2010: th e continental shelf project of the kingdom of denmark – status at the beginning of 2010. geological survey of denmark and greenland bulletin 20, 51–54. marcussen, c., christiansen, f.g., dahl-jensen, t., heinesen, m., lomholt, s., møller, j.j. & sørensen, k. 2004: exploring for extended continental shelf claims off greenland and the faroe islands – geological perspectives. geological survey of denmark and greenland bulletin 4, 61–64. recommendations of the commission on the limits of the continental shelf in regard to the partial submission made by the government of denmark together with the government of the faroes in respect of the continental shelf north of the faroe islands on 29 april 2009. adopted by the commission, with amendments, on 12 march 2014. http://www. un.org/depts/los/clcs_new/submissions_ fi les/dnk28_09/2014_03_14_ scdnk_rec_com_20140521.pdf weigelt, e. & jokat, w. 2001: peculiarities of roughness and thickness of oceanic crust in the eurasian basin, arctic ocean. geophysical journal international 145, 505–516. 0.0 0.5 1.0 1.5 2.0 d ep th b el o w s ea flo o r (k m ) 1.5 2.0 2.5 3.0 3.5 4.0 velocity (km/s) fig. 4. sediment velocities in the amundsen basin obtained from ray-tracing of the sonobuoy recordings from the three geus lomrog expeditions. green circles show individual data points; the red line indicates an approximation obtained from reduced major axis regression techniques. authors' address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: cma@geus.dk geological survey of denmark and greenland bulletin 41, 2018, 99-104 99 photogrammetry is a classical remote sensing technique dating back to the 19th century that allows geologists to make three-dimensional observations in two-dimensional images using human stereopsis. pioneering work in the 1980s and 1990s (dueholm 1992) combined the use of vertical (nadirlooking) aerial photographs with oblique stereo images from handheld small-frame cameras into so-called multi-model photogrammetry. this was a huge technological step forward that made it possible to map, in three dimensions, steep terrain that would otherwise be inaccessible or poorly resolved in conventional nadir-looking imagery. the development was fundamental to the mapping and investigation of e.g. the nuussuaq basin (pedersen et al. 2006). digital photogrammetry, the all-digital version of multi-model photogrammetry, is nowadays an efficient and powerful geological tool that is used by the photogeological laboratory at the geological survey of denmark and greenland (geus) to address geological problems in a range of projects from 3d mapping to image-based surface reconstruction and orthophoto production. here we present an updated description (complementary to dueholm 1992) of the analytical procedures in the typical digital workflow used in current 3dmapping projects at geus. analytical procedures for 3d mapping at the photogeological laboratory of the geological survey of denmark and greenland erik vest sørensen and mads dueholm fig. 1. during field work new stereo imagery is normally collected from helicopters or boats, but could also be collected from smaller fixed-wing aircraft, drones or while walking. © 2018 geus. geological survey of denmark and greenland bulletin 41, 99–104. open access: www.geus.dk/bulletin http://www.geus.dk/bulletin 100100 multi-model photogrammetry in its present form is essentially a technique that allows geologists to combine stereo images of different origin e.g. from satellite, aerial or handheld cameras, and with different resolutions and viewing angles in their geological interpretations (fig.1), using a digital photogrammetric workstation. examples are plentiful with scales ranging from metres (vosgerau et al. 2010, 2016) to kilometres (svennevig et al. 2015; sørensen & guarnieri 2018, this volume) or even hundreds of kilometres (sørensen et al. 2017). the strength of the methodology lies in the ability to combine stereo images with different viewing angles. regional geological structures are e.g. typically well resolved in nadir-looking aerial or satellite images, while steep cliffs are better resolved from closer range images acquired perpendicularly to the slope of the outcrop. essentially, digital photogrammetry allows the user to map and quantify in three dimensions whatever can be seen in the stereo images across different scales and resolutions. the digital photogrammetric workstation described here is the modern equivalent to the analytical setup previously used at geus (hougaard et al. 1991) and at the institute of surveying and photogrammetry, technical university of denmark (dueholm 1992). the workstation consists of a windows-based computer and a split-screen 3d monitor system (fig. 2). technically, the monitor system displays an image of an object in one screen and an overlapping image of the same object but from a slightly different position on a second screen. a beam splitter mirror splits the polarisation direction of the two screens into separate horizontal and vertical directions. polarising glasses used by the viewer then filter the signals so that the top screen is solely presented to the right eye and the bottom screen is solely presented to the left eye. this allows the human visual system to merge the two images, whereby a stereoscopic model is created in front of the user. the central part of the workstation is the photogrammetric software, which is essentially a computer-controlled set of algorithms that controls the viewing of the stereoscopic model as well as the collection and manipulation of threedimensional data within the model. this allows for seamless movement and data capture between different stereoscopic models regardless of scale, origin and viewing angle, which highly increases the speed and efficiency of geological mapping, especially in steep and inaccessible terrains. at present, the photogeological laboratory uses two commercial photogrammetric 3d-mapping software solutions (socet gxp from bae systems and 3d stereo blend) that are complementary to each other in terms of technical capabilities. in the following, we describe the typical workflow used with 3d stereo blend: 1) data acquisition during field work, 2) data preparation of the images, 3) data interpretation (fig. 3). this description, however, should not be viewed as a complete manual to the software. 3d stereo blend is developed by anchor lab in close collaboration with geus’ photogeological laboratory. the software is optimised to 3d-mapping and structural interpretation using oblique stereo images collected with calibrated hand-held digital cameras. it is an essential part of the overall strategy of the photogeological laboratory to increase the efficiency and usability of digital photogrammetry from data acquisition to end-product, so that the method becomes a geological tool routinely used by geologists also without prior expert knowledge. fig. 2. the geologist at work in the photogeological laboratory. the stereoscopic model is displayed on a 3d monitor system that is well suited for full-day working. the stereoscopic model can be observed simultaneously by several viewers; this makes it easy to illustrate and discuss geological observations and ideas, which is beneficial for the geological interpretation. 101 data acquisition – field photography field work provides the geologist with important first order observations of the bedrock. working in remote and mountainous areas such as greenland is often challenging because of inaccessible outcrops (steep cliffs) and time constraints (short field seasons). the result is spatially scattered outcrop observations that can be difficult to correlate or relate to overall regional structures. to overcome this problem, stereo images, i.e. strips of overlapping images taken from different positions and covering the geological outcrops of interest are collected on a routine basis using hand-held digital cameras, commonly deployed from helicopters (fig. 4). the simplicity of using a hand-held digital camera makes the data acquisition extremely mobile and fast. furthermore, the quality, resolution and storage capacity of modern digital cameras have led to a huge increase in efficiency and capacity during data acquisition compared to the earlier days of 3d mapping. as an example, overview images of a small outcrop were collected within minutes using a helicopter, while on a more regional scale, a 100 km cliff section was photographed within less than one hour (sørensen et al. 2015a). depending on the logistic setup, the images could equally well be collected by other means (fig. 1), as long as sufficient overlap between the images is ensured. as a rule of thumb, 60–80% overlap is needed to obtain continuous overlap and good stereoscopic measurement accuracy (dueholm 1992). however, it is now recommended that images are taken with up to 90% overlap because this results in a more successful automatic generation of common points (so-called tie points) between different images. it will also, at a later stage, make it possible to use the images for surface reconstruction using dense image matching routines (sørensen et al. 2015b). the cameras are full-frame, digital single-lens reflex with high-quality 35 mm prime lenses that are fixed and focused at infinity. however, essentially any camera can be used as long as the camera parameters (lens distortion, focal length and principal point) of the camera can be modelled. the cameras are calibrated prior to field work using a test field consisting of a steel grid with c. 100 points. we recommend that the images are acquired in the raw image format of the camera and that the location of the camera is registered with global navigation satellite systems (gnss). depending on the requirements for absolute accuracy, different gnss equipment can be used, from simple geotagging devices to more advanced differential gnss systems. data preparation setting up the images essentially consists of two steps. first, a relative model is constructed by identifying common points (tie points) between overlapping images. secondly, this model is transformed into absolute, ‘real-world’ coordinates by combining camera location data and control points. relative orientation the relative connectivity between images in object space is established by automatic tie point measurement using ‘structure from motion’ (sfm) image-matching algorithms (lowe 2004; snavely et al. 2008), with possible manual editing or addition of tie points by stereoscopic measurement within 3d stereo blend. the commercial software photoscan professional (ps) from agisoft is used for the image matching, but other software solutions can also be used. the raw imagematches (tie points) are subsequently exported to 3d stereo blend. 3d stereo blend uses the imported tie point file as a block definition and set-up file as well as for preliminary triangulation. the file is typically thinned when imported using an area-based thinning technique to give an even distribution of tie points across the images. the next step is to make a preliminary so-called bundle adjustment or triangulation, including error detection and elimination of erroneous tie points from the imported data. the images can subsequently be viewed in 3d with the orientation of that exported from ps. if positional data, such as e.g. gnss camera positions, are included in the export from ps, the stereofield acqusition using calibrated digital cameras automatic image matching using structure from motion (photoscan) traditional photogrammetric bundle adjustment (3d stereo blend) archive stereo imagery surface reconstruction (point cloud/dem) (sure & photoscan) 3d mapping (3d stereo blend) data interpretation data preparation data acquisition fig. 3. schematic flow diagram summarising the typical 3d-mapping workflow from data acquisition to photogrammetric data preparation and interpretation. 102102 scopic model will be placed in absolute coordinates within 3d stereo blend. depending on the project requirements, one could move directly on to the 3d-mapping stage for rapid results or if absolute accuracy is of minor importance. however, requirements for absolute accuracy often mean that additional control data must be added. absolute orientation the absolute orientation relates the photogrammetric models in object space to absolute coordinates through a proper bundle adjustment process, whereby the stereoscopic model gets the correct scale and levelling. when solving the bundle adjustment, all provided control information is weighted according to an a priori estimated error. different sources of control data can be used in this process, including surveyed control points, pass-points, camera gnss-data, planar levelling points and distances. surveyed ground-control points generally give the best positional accuracy. however, considering the regional scale of many 3d-mapping projects, this approach is often not viable from a practical point of view. instead, pass-points from already aerotriangulated aerial photographs can be used as control-point source, which eliminates the need for ground-control-point collection during field work. this is done by identifying common points between for example a set of vertical aerial photographs and local oblique-view images. this process is sensitive due to the different perspectives of the image data sets as well as the different resolutions, and takes some practical experience to carry out. however, workflows implemented within 3d stereo blend have significantly improved the efficiency of the identification process. gnss data collected with the camera yield the position of the camera at the time of acquisition that is important for setting up the images. simple geotagging equipment typically yields a camera position accuracy of around 5–10 m. however, more advanced differential gnss set-ups could result in a positional accuracy at the sub-metre level, which might minimise or completely eliminate the need for surveyed groundcontrol points or measured passpoints. the overall time consumption in the preparation of the images has significantly decreased compared to the early days of analytical 3d mapping. this is largely a consequence of computer hardware and digital camera development, but also of software improvements including better image-matching algorithms and improved photogrammetric workflows in the 3d stereo blend software. consequently, small blocks of stereo images can be prepared for geological interpretation within a day. this opens up for e.g. using digital photogrammetry as an active tool during field work. stereoscopic image interpretation – 3d mapping once the images are properly prepared, or if archive data are available, the geologist can commence the geological photointerpretation in three dimensions. this is done by tracing geological features of interest in the stereoscopic models. the change between neighbouring stereo models takes place automatically, so that important horizons can be traced seamlessly for kilometres. several views can be opened simultaneously, whereby an outcrop can be seen from different perspectives and at different resolutions, which is important for the geological interpretation. the different views can furfig. 4. stereo images collected from helicopter using a hand-held digital camera. this method makes the data acquisition very mobile and fast. the location of the camera is registered with global navigation satellite systems. photo: jonas petersen. 103 thermore be linked, so that a movement in one window is also updated in the linked window. changing the displayed stereo block is either done by selecting from a pull-down window or by selecting the cameras in perspective view interactively. all images from the cameras within a project can be shown in the perspective view. this makes it easy to manage and move around in large regional data sets with thousands of images. the outcome of the drawing is a number of vectorised lines in 3d with many nodes (so-called polylines) superimposed on the stereoscopic model (fig. 2). the polylines can be labelled and grouped according to the user’s need, and subsequently exported to gis packages for gis analysis or to 3d-modelling software. the software automatically registers the original stereo model in which each node of a given polyline was drawn. this makes it possible to automatically adjust the images if their orientation is changed at some point. this is helpful e.g. if an initial interpretation of the images is done on a preliminary set-up of the images without proper ground control. a central part of the 3d stereo blend software is a set of structural tools that enables the viewer to evaluate structural parameters such as strike and dip of bedding, plunge and direction of fold axes and stratigraphic thickness of beds. this allows the geologist for instance to populate a geological map with many structural observations from areas that could not be visited in the field, to correlate bedding from one side of a fjord to the other or to measure true thicknesses. the structural tool-set consists of selected routines from the geoprogram software (dueholm & coe 1989) and works by fitting planes by least-squares adjustment of captured data points. of special importance is the possibility to project captured data into geological sections which can have arbitrary orientations and also be inclined to better reflect e.g. true thicknesses. in summary, with the present set-up, 3d mapping has become much more effective and user-friendly than previously, which is largely due to the improved photogrammetric workflows in the 3d stereo blend software. for instance, it is now possible for the untrained geologist to engage in 3d mapping with only 1–2 days of training. resolution and accuracy the resolution of the stereoscopic models depends on the distance between the camera and outcrop, the camera focal length and the pixel size (pixel pitch) of the sensor of the digital camera. as an example, photographing an outcrop at a distance of 100 m gives a ground sampling distance (gsd) of 14 mm using e.g. a 36 megapixel nikon d800e camera with a calibrated sensor pixel pitch of 4.89 µm and a focal length of 36 mm, or a scale of c. 1:3000, while increasing the distance to 1000 m will give images with gsd of 0.14 m (scale 1:30 000). the photogrammetric or geometric accuracy in the image plane on the ground (the x and y axes perpendicular to the direction of view) relates to the distance between the camera and outcrop, the camera focal length and how accurately a point can be determined in the stereoscopic model. the latter is a function of the accuracy of the triangulation, the camera calibration and on how well the user can place a point in the stereoscopic model. in addition, the photogrammetric accuracy in depth (the z axis, in the direction of view) also depends on the ratio between the distance between the camera and the object and the distance between the camera stations (also called the baseline). a typical value for the point determination is around one pixel, which leads to an accuracy of c. 14 mm in the image plane, while the accuracy in depth is c. 35 mm with a distance to the outcrop of 100 m and a baseline of 40 m (corresponding to 60% overlap). if the camerato-object distance is increased to 1000 m, the accuracy in the image plane decreases to 0.14 m, while the accuracy in depth will decrease to 0.35 m, assuming that the baseline is increased to 400 m to maintain a 60% image overlap. however, if the baseline remains 40 m long there is a significant decrease in the depth accuracy to c. 3.5 m. this illustrates the importance of having an appropriate image overlap. the absolute accuracy of the stereoscopic model relates to that of the control source. in greenland, pass-points are commonly acquired from the monochrome vertical aerial photographs on a scale of 1:150 000 that cover most ice-free areas. typical achievable accuracies on the point transfer from these photographs is around 3–5 m. using high-precision differential gnns setup it should be possible to obtain accuracies of less than 1 m, whereby the absolute accuracy approaches the photogrammetric accuracy. although the absolute accuracy generally exceeds that of the photogrammetric accuracy, the relative accuracy between models remains equal to the photogrammetric accuracy, because the ground-control data is weighted during the bundle adjustment. this means that when calculating e.g. thicknesses or structural parameters such as strike and dip, the accuracy is determined from the photogrammetric accuracy. in practice, this is all handled automatically within the 3d stereo blend software. other derived products in addition to the 3d geological mapping workflow, the technical development of automatic multi-viewstereo-matching routines (rothermel et al. 2012) has facilitated the extraction of digital outcrop models from stereo imagery (sørensen 104104 et al. 2015b). the digital outcrop model is a 3d representation of the outcrop surface. this type of routine utilises the redundancy of high image overlap to produce a set of highresolution data points in space (so-called point clouds) that need little manual editing. although a high image overlap (i.e. a small baseline) reduces the precision of individually matched pixels, this is compensated by determining the same point in multiple images, which leads to effective automatic elimination of erroneous points. the point cloud is used e.g. for visualisation purposes where it can be integrated with the results of the 3d geological mapping, but it can also be further processed into 3d mesh representations of the terrain, or production of digital terrain models and orthophotos. geus’ photogeological laboratory is currently using a suite of software solutions for terrain extraction. software such as photoscan professional from agisoft and sure from nframes is typically used in the digital outcrop model workflow, while more conventional aerial and satellite imagery is processed using socet gxp. summary the all-digital version of multi-model photogrammetry, now referred to as digital photogrammetry or just 3d mapping, has brought the geological outcrop into geus’ photogeological laboratory. the recent increases in efficiency all the way from data acquisition to the geological interpretation makes 3d mapping an attractive geological tool available to the geologist. with a digital photogrammetric workstation, the users can view, map and explore any geological feature in three dimensions, following the principle that whatever can be seen in the images can also be mapped and quantified in 3d. acknowledgments the authors would like to thank max strunck and birgir óskarsson for reviewing the manuscript as well as asger ken pedersen and lotte melchior larsen for helpful comments and suggestions. references dueholm, k.s. & coe, j.a. 1989: geoprogram. program for geological photogrammetry. the compass 66, 59–64. dueholm, k.s. 1992: geologic photogrammetry using standard smallframe cameras. in: dueholm, k.s. & pedersen, a.k. (eds): geological analysis and mapping using multi-model photogrammetry. rapport grønlands geologiske undersøgelse 156, 7–17. hougaard, g., jepsen, h.f. & neve, j.k. 1991: ggu’s photogeological laboratory: aerial photogrammetry – a valuable geological mapping tool in greenland. grønlands geologiske undersøgelse rapport 152, 32–35. lowe, d.g. 2004: distinctive image features from scale-invariant key points. international journal of computer vision 60, 91–110. pedersen, a.k., larsen, l.m., pedersen, g.k. & dueholm, k.s. 2006: five slices through the nuussuaq basin, west greenland. geological survey of denmark and greenland bulletin 10, 53–56. rothermel m., wenzel, k., fritsch, d. & haala, n. 2012: sure: photogrammetric surface reconstruction from imagery. in: proceedings, lc3d workshop, berlin, 4–5 december 2012, 9 pp. snavely, n., seitz, s. & szeliski, r. 2008: modeling the world from internet photo collections. international journal of computer vision 80, 189–210. sørensen, e.v., pedersen, a.k., garcia-sellés, d. & strunck, m.n. 2015a: point cloud from oblique stereo-imagery: an outcrop case study across scales and accessibility. european journal of remote sensing 48, 593– 614. sørensen, e.v., bjerager, m. & citterio, m. 2015b: digital models based on images taken with handheld cameras – examples on land, from the sea and on ice. geological survey of denmark and greenland bulletin 33, 73–76. sørensen, e.v., baker, n.g. & guarnieri, p. 2017: three years of photographing – extreme 3d mapping. abstract, grsg 28th international annual conference – applied geological remote sensing. https:// www.grsg.org.uk/wp-content/uploads/2017/12/grsg-agm-and-conference-abstract-book-2017.pdf sørensen, e. v. & guarnieri, p. 2018: remote geological mapping using 3d photogrammetry: an example from karrat, west greenland. geological survey of denmark and greenland bulletin 41, 63–66 (this volume). svennevig, k., guarnieri, p. & stemmerik, l. 2015: from oblique photogrammetry to a 3d model – structural modeling of kilen, eastern north greenland. computers & geosciences 83, 120–126. vosgerau, h., guarnieri, p, weibel, r., larsen, m., dennehy, c., sørensen, e.v. & knudsen, c. 2010: study of a palaeogene interbasaltic sedimentary unit in southern east greenland: from 3-d photogeology to micropetrography. geological survey of denmark and greenland bulletin 20, 75–78. vosgerau, h., passey, s. r., svennevig, k., strunck, m. n. & jolley, d.w. 2016: reservoir architectures of interlava systems: a 3d-photogrammetrical study of eocene cliff sections, faroe islands. geological society, london, special publications 436, 55–73. authors’s addresses e.v.s., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: evs@geus.dk. m.d., anchor lab k/s, h.c. andersens boulevard 37, 5. mf, dk-1553 copenhagen v, denmark. https://www.grsg.org.uk/wp-content/uploads/2017/12/grsg-agm-and-conference-abstract-book-2017.pdf https://www.grsg.org.uk/wp-content/uploads/2017/12/grsg-agm-and-conference-abstract-book-2017.pdf https://www.grsg.org.uk/wp-content/uploads/2017/12/grsg-agm-and-conference-abstract-book-2017.pdf mailto:evs@geus.dk geological survey of denmark and greenland bulletin 31, 2014, 75-78 75 earthquake swarms in greenland tine b. larsen, peter h. voss, trine dahl-jensen and hans peter rasmussen two earthquake swarms have been detected in greenland. one occurred on the island of disko in august 2010, the other one was active from january 2008 to june 2009 near the south-east greenland coast c. 200 km south of tasiilaq. an earthquake swarm is defined as a series of earthquakes of similar magnitude located within a small area. the magnitude of the largest earthquakes in a swarm is typically less than 4 (ma & eaton 2009). swarm activity is distinctly different from the more common mainshock–aftershock activity, which is characterised by one large earthquake (mainshock) followed by a series of smaller aftershocks. earthquake swarms mainly occur in areas with tectonic and/or volcanic activity (stykes 1970), but intraplate swarms are also found in otherwise stable environments (gregersen 1979; atakan et al. 1994; uski et al. 2006; ma & eaton 2009). geological boundaries and old fault zones appear to be a common setting for intraplate earthquake swarms. earthquake swarms have previously been detected in north and north-east greenland (gregersen 1979) at a time when the seismograph coverage was very sparse. it was concluded that the earthquake swarms were caused by tectonic stresses in and around old sedimentary basins near the continental margin. in this study we take advantage of the recently improved network of digital broadband seismographs in greenland (dahl-jensen et al. 2010). the shorter distance between seismograph stations and the high quality digital data enable us to better detect small earthquakes. we have focused on the last decade during which the digital network has gradually been established (fig. 1). we revisited the two regions where swarms have previously been detected, in north greenland and in north-east greenland (gregersen 1979). searching the geus earthquake database for recent earthquakes showed nothing in the area of the north greenland swarm. the swarm described by gregersen (1979) was primarily detected using data from a canadian station (ale) and a temporary station on the ice sheet, supplemented with data from station nord (nor) for the largest earthquakes in the swarm. this particular region has not experienced an increase in station coverage in recent years. it is therefore impossible to distinguish whether the lack of recent detections is due to a lack of activity in that area. the location of the north-east greenland earthquake swarm has also been searched without finding any new swarm activity in the last decade. only two small earthquakes have been detected in the area, one in 2008 and another in 2011. in this area the station coverage has improved since 1974, so the lack of detections reflects that the area is currently seismically quiet. two new localities with earthquake swarms were uncovered in this study (fig. 1). both are areas with regular seismic activity. one swarm is located on the island of disko near the disko gneiss ridge (fig. 2) where a fault zone runs n–s through the island (chalmers et al. 1999). this is a very distinct swarm with its main activity concentrated on two days © 2014 geus. geological survey of denmark and greenland bulletin 31, 75–78. open access: www.geus.dk/publications/bull dag sco sfjd nrs nuuk sumg angg (in tasiilaq) dy2g dbg soeg norale iluli kullo tuleg isog ivi nuug icesg upnv fig. 2 fig. 3 250 km 60°n 60° 50° 40° 30° 20° 10° 0° 40°w 80°n 70° disko umiivik archaean volcanic rocks palaeozoic fold belts sedimentary basins and intrusions reworked archaean palaeoproterozoic precambrian shield neem fig. 1. geological map of greenland with seismograph stations marked by triangles. the locations of the swarms found by gregersen (1979) are marked with red circles. the locations of the swarms presented in this study are marked with black boxes. 7676 in late august 2010, followed by slightly increased seismicity during the following two months. the other swarm is located in a remote area near the south-east greenland coast, c. 200 km south of tasiilaq (fig. 3). this swarm is less distinct with two main pulses of activity in january 2008 and spring 2009. this area had increased seismicity for a year and a half, before it tapered off. the objective of this paper is to describe the two newly discovered earthquake swarms and how they relate to the local geology. data and analysis seismographs have been continuously operating in greenland since 1927 (hjelme 1996). originally the instruments were large, difficult to install and labour intensive to maintain. it was only possible to operate a few stations in greenland. this changed with the development of digital instruments, and during the last decade a network of 20 digital broadband seismographs with real-time data transmission has been installed in greenland (fig. 1). this massive undertaking is a result of the international collaboration in the greenland ice sheet monitoring network (glisn) project with funding from around the world (dahl-jensen et al. 2010). the latest seismological station was installed in upernavik in 2013 by the korean polar research institute (see upnv fig. 1). data from the real-time network of seismographs are processed daily at geus using the seisan analysis software (ottemöller et al. 2013). earthquake phases are manually identified and combined to form earthquake locations when possible. locations and phase readings are kept in the geus database for future reference. this includes phase readings not associated with an earthquake location. all the raw waveform data are also stored. the new earthquake swarms were discovered by performing a search in the database around locations, where clustering of seismicity was observed on a map of all earthquakes in greenland. the search results were then examined for possible clustering of events in time. two earthquake clusters fulfilled the criteria defining an earthquake swarm. despite the improved station coverage, it is still challenging to determine the focal mechanism of most earthquakes in greenland. particularly the small swarm earthquakes that are only registered on one to four seismographs, which is insufficient for reliable focal mechanism calculations. earthquake swarm at the island of disko disko and the surrounding areas are frequently shaken by small earthquakes, some large enough to be felt by the locals. disko is also known for its warm springs formed where circulating surface water penetrates into the subsurface through deep cracks and is heated by the higher temperatures at depth. the flow of a warm spring is sometimes changed by a small earthquake, indicating that the earthquake sources are shallow. there are thousands of springs with constant temperature on disko. the temperatures in individual springs vary between 1°c and 18.5°c (kristensen 2006). the number of earthquakes detected near the southern end of the disko gneiss ridge per year for 2006–2013 varies between one and ten (fig 4a). however, 2010 stands out with 27 detected earthquakes. the earthquake activity was particularly intense on 22 and 23 august (fig 4b). the earthquakes ranged in magnitude from 1.9 to 3.2 with most lower tertiary basalts upper cretaceous and lower precambrian basement tertiary sediments fault10 km 69°30´ storbræen disko fjord 55°w fig. 2. the calculated epicentres for the swarm earthquakes on disko, west greenland. the disko gneiss ridge runs along the fault line on the map. reworked archaean archaean 64°30´ 10 km umiivik 42°w 64°n 64°30´n 64°n 40°w 40°w fig. 3. the calculated epicentres for the swam earthquakes near umiivik, on the south-east greenland coast. 77 clustering in the magnitude range between 2.2 and 2.8. the earthquake epicentres were spread out over a small area with the largest distance between two events being about 30 km. this distance is within the uncertainty on the locations. the uncertainty on the epicentres is in the order of 10 km in n–s direction and at least 50 km in e–w direction. this is caused by the geometry of the network relative to the events with seismographs primarily located north and south of the epicentres. as fig. 2 shows, the scatter in the epicentres was indeed largest in e–w direction. the closest seismograph where all earthquakes were registered is located in ilulissat (iluli) 110–120 km away. the earthquakes were also registered in nuugaatsiaq (nuug) c. 220 km to the north, in kangerlussuaq (sfjd) 320 km to the south and at the summit ice camp (sumg) 650 km towards the north-east. a few earthquakes were registered on a total of eight seismographs in greenland, but most of them were registered in just four locations. the depth of the earthquakes is poorly determined, but appears to be shallow. a major challenge in determining the depth is the limited knowledge of the crustal velocity structure. the depth to moho is known from receiver function studies (dahl-jensen et al. 2003), and the crustal density structure in the disko area has been modelled from seismic and potential field data (chalmers et al. 1999). relocating the earthquakes with a fixed depth of 5 km does not reduce the distance between the epicentres significantly. however, the variation in time difference between the p-wave arrival and the s-wave arrival was less than 1.5 sec. indicating a small source region. the north–south-oriented fault along the disko gneiss ridge is a known zone of weakness, and the area is characterised by a moderate earthquake activity with no seasonal variation. in se finland earthquake swarms are observed to be related to the intrusion of water and gas into fractures working in concert with the local stress field (uski et al. 2006). a similar mechanism is likely on disko. as warm water is circulating in deep cracks all year round a seasonal variation is not expected. some of the warm springs on disko are radioactive (kristensen 2006) containing radon from the basement. radon is also observed near earthquake swarms in finland, and radon gas is believed to play a minor role in triggering earthquake swarms in se finland (uski et al. 2006). earthquake swarm in south-east greenland the east coast of greenland is frequently shaken by small earthquakes. in tasiilaq earthquakes are felt almost every year. many small earthquakes are registered farther down the coast, but it is not known if any of these events can be felt as the area is not populated. earthquake swarm activity was detected in january 2008 and in the spring of 2009 just off the coast near umiivik. the earthquakes in the swarm range in magnitude from 1.8 to 3.8. in the years before and after the swarm period, the level of earthquake activity in the area range between 12 and 21 earthquakes per year (fig. 5a). however, in 2008 and 2009 respectively 71 and 114 earthquakes were registered. the earthquakes occurred primarily in january and april 2008 and again from january to july 2009 (fig. 5b). the uncertainty on most of the epicentres is several hundred kilometres in both the n–s and e–w directions. this is a consequence of the earthquakes being detected primarily by only two seismographs (four phases in total) in kangerlussuaq (sfjd) and at summit (sumg) 550 km and 1000 km from the epicentres. the distance between the seismographs and the epicentres is well-determined, but the direction is poorly constrained. after improvements had been made to the seismograph network, a larger earthquake with a magnitude of 3.6 occurred in the area on 13 october 2009. this earthquake was well-recorded at four locations: narsarsuaq (nrs), kangerlussuaq (sfjd), ilulissat (iluli) and summit (sumg) with an uncertainty of less than 10 km in n–s direction and less than 20 km in e–w direction. this earthfig. 5. a: number of earthquakes per year for the period 2006–2013 around the location of the earthquake swarm at the se greenland coast. b: number of earthquakes per month from january 2008 to july 2009 in the same area. 0 1 2 3 4 5 20 /0 820 10 21 /0 820 10 22 /0 820 10 23 /0 820 10 24 /0 820 10 25 /0 820 10 26 /0 820 10 27 /0 820 10 28 /0 820 10 29 /0 820 10 30 /0 820 10 0 5 10 15 20 25 2006 2008 2010 2012 year date n um be r o f e ar th qu ak es n um be r o f e ar th qu ak esa b disko fig. 4. a: number of earthquakes per year for the period 2006–2013 around the location of the earthquake swarm on disko. b: number of earthquakes per day from 20 to 30 august 2010 in the same area. jan janm ar m ay ju l se p n ov m ar m ay ju l2006 2008 2010 2012 year month (2008 and 2009) n um be r o f e ar th qu ak es n um be r o f e ar th qu ak es 0 20 40 60 80 100 0 5 10 15 20 a bse greenland 7878 quake is located near the centre of the cloud of epicentres, and it can be used to evaluate the relative locations of the swarm earthquakes. the waveforms for the small earthquakes are very noisy, making a direct comparison difficult. instead we examine the variation in time difference between the p-wave and swave arrival for individual earthquakes. despite the epicentres being scattered over almost 100 km in n–s direction and c. 50 km in e–w direction, the variation in s–p time (i.e. the distance to the earthquake) at each seismograph station fluctuated only in rare cases by more than 1.5 sec. relative to the 13 october 2009 earthquake. the large scatter in epicentres is probably an artefact of the focal depth being unconstrained. the umiivik area marks a major geological transition between the relatively unreworked archaean rocks to the south and rocks highly deformed in the proterozoic to the north (henriksen et al. 2009). offshore the existence of a failed rift arm was proposed by hopper et al. (1998). the earthquake swarm was thus located close to two old geological boundaries. this is in good accordance with the general observations by stykes (1978) that intraplate earthquakes indeed occur in old zones of weakness. conclusion as shown by this initial search for earthquake swarms, the improvements to the seismograph network in greenland enable us to better uncover and analyse swarms in the future. the swarms presented in this paper could not have been detected a decade ago. intraplate earthquake swarms can contribute to the understanding of geological processes currently at work near old geological boundaries in remote areas. both swarms presented here were active near old geological boundaries. the swarm on disko is located at the southern end of the disko gneiss ridge where a large fault runs through the island. circulating water in deep cracks may possibly play a role. the se greenland swarm is located close to a major geological boundary between reworked and unreworked archaean rocks. furthermore, the swarm is close to the proposed location of a failed rift arm in the atlantic. finding earthquake swarms close to old, major, geological boundaries is well in accordance with the literature. the earthquake swarms presented in this paper are not associated with known sedimentary basins such as the swarms described in gregersen (1979). this indicates a diversity of swarm sources in greenland. further work will be carried out on the earthquake swarms in greenland and data from temporary and canadian seismographs will be included in the analysis. acknowledgement the data were collected and distributed by the glisn project and its members, www.glisn.info. references atakan, k., lindblom, c.d. & havskov, j. 1994: earthquake swarm in steigen, northern norway: an unusual example of intraplate seismicity. terra nova 6, 180–194. chalmers, j.a., pulvertaft, t.c.r., marcussen, c. & pedersen, a.k. 1999: new insight into the structure of the nuussuaq basin, central west greenland. marine and petroleum geology 16, 197–224. dahl-jensen, t. et al. 2003: depth to moho in greenland: receiver function analysis suggests two proterozoic blocks in greenland. earth and planetary science letters 205, 379–393. dahl-jensen, t., larsen, t.b., voss, p.h. & the glisn group 2010: greenland ice sheet monitoring network (glisn): a seismological approach. geological survey of denmark and greenland bulletin 20, 55–58. gregersen, s. 1979: intraplate earthquake swarms in greenland and adjacent continental regions. nature 281, 661–662. henriksen, n., higgins, a.k., kalsbeek, f. & pulvertaft, t.c.r. 2009: greenland from archaean to quaternary. descriptive text to the geological map of greenland 1:2 500 000. 2nd edition. geology of greenland survey bulletin 18, 126 pp. hjelme, j. 1996: history of seismological stations in denmark and greenland. in: wahlström, r. (ed.): seismograph recording in sweden, norway – with arctic regions, denmark – with greenland, and finland. proceedings from the uppsala wiechert jubilee seminar, 49–57. uppsala: seismological department, uppsala university, sweden. hopper, j.r., lizarralde, d. & larsen, h.c. 1998: seismic investigations offshore south-east greenland. geology of greenland survey bulletin 180, 145–151. kristensen, r.m. 2006: de varme kilder. in: bruun, l. et al. (eds): arktisk station 1906–2006, 310–315. copenhagen: rhodos. ma, s. & eaton, d.w. 2009: anatomy of a small earthquake swarm in southern ontario, canada. seismological research letters 80, 214– 223, http://dx.doi.org/10.1785/gssrl.80.2.214 ottemöller, l., voss, p.h. & havskov, j. 2011: seisan earthquake analysis software for windows, solaris, linux and macosx. university of bergen. stykes, l.r. 1970: earthquake swarms and sea-floor spreading. journal of geophysical research 75, 6598–6611. stykes, l.r. 1978: intraplate seismicity, reactivation of pre-existing zones of weakness, alkaline magmatism, and other tectonism postdating continental fragmentation. reviews of geophysics and space physics 16, 621–688. uski, m., tiira, t., korja, a. & elo, s. 2006: the 2003 earthquake swarm in anjalankoski, south-eastern finland. tectonophysics 422, 55–69. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark; e-mail: tbl@geus.dk geological survey of denmark and greenland bulletin 38, 2017, 17-20 17 the water framework directive (wfd) of the european union prescribes “good ecological status” of all waters. in terms of nitrate this means, among other things, to avoid eutrophication and achieve a good ecological balance in surface water systems for the benefit of the groundwater dependent flora and fauna (hinsby et al. 2012). in denmark, the nitrate load to estuaries has been nearly halved since the first national action plan was implemented in the mid-1980s, but further abatements are required in many areas to fulfil the wfd. new approaches to regulate nitrate use are needed with measures targeted to the areas where most effect is obtained, and this is recognised at political level. recent legislation allows farmers to increase nitrate application, but should at the same time introduce new mitigation measures and a more targeted approach to regulation. therefore the physical system, i.e. the geological framework and topography, of the catchment has to be understood (winter 1999). previous studies have shown that in hydrological catchments with high geological variability, sampling of groundwater in riparian zones, the stream water itself and water in the stream bed can help to identify near-stream areas with specific nitrate problems. detailed studies are, however, not feasible in all catchments, and development of representative typologies to guide an optimal location of mitigation measures in the catchment is thus needed. the present study is a detailed characterisation of nitrate transport and reduction in the groundwater–stream system in the river hagens møllebæk catchment for this purpose. study area and hydrogeological setting the hagens møllebæk catchment is located west of skive and discharges into skive fjord (fig. 1). the hydrological catchment covers 27.2 km2. the area is relatively flat with elevations between 1–52 m above sea level (a.s.l.) and slopes near the streams. quaternary clayey tills dominate in the west and south and glacial meltwater sand and gravel occur in the east. holocene sand, clay and organic deposits overlie the quaternary deposits in the stream-valley system and near the outlet of hagens møllebæk. in the southern part of the catchment, oligocene clay with very low permeability is found underneath a few metres of clayey till . the primary land use is agriculture, which covers more than 80% of the area. the soil types are clay (50%), sand (47%) and organic sediments (3%). nitrate transport pathways in riparian zones of the hagens møllebæk catchment, northern denmark bertel nilsson, anker lajer højberg and per jensen •••••• • • • • • • • • • • 9 8 7 6 5 4 2 1 14 13 12 10 3 15 11• intensive • extensive stream hagens møllebæk catchment 52 m 1m 2 km monitoring station elevation skive fjord hagens m ølleb æk hagens mølle jebjergdenmark sweden 50 km jylland skive ab fig. 1. a: topographical map of the hagens møllebæk catchment with locations of nitrate-measuring stations of riparian zones, drain pipe outlets, stream and stream bed. b: index map with position of study area. fig. 2. conceptual geological profile along hagens møllebæk with locations of numbered nitrate-measuring stations along the main stream channel. 1 248 12 13 1415 0 1 2 3 4 5 6 7 8 km -10 0 10 20 al tit ud e (m ) jebjerg hagens mølle holocene sand and clay late glacial sand clayey till meltwater sand and gravel oligocene clay © 2017 geus. geological survey of denmark and greenland bulletin 38, 17–20. open access: www.geus.dk/publications/bull 1818 conceptual geological model a conceptual geological model is presented in fig. 2 by a profile along the main stream of hagens møllebæk from jebjerg in the north (near station 15, fig. 1a) to hagens mølle (station 1) near the stream outlet into skive fjord. the model is based on existing geological data from the jupiter database and a quaternary soil map on a scale of 1:200 000 (both found at www.geus.dk). from the conceptual model, a general understanding of the contact and interaction between the stream and the underlying aquifers is established on catchment scale. between stations 4 and 5 the stream bed overlies poorly permeable clayey tills or oligocene clay with little or no expected water exchange. conversely, exchange of water and nitrate between the stream and the underlying aquifer is more likely where the stream bed is located directly on top of sandy aquifers. regional contact between aquifers and riparian zones the local geological and hydrogeological conditions in the riparian zone of the hagens møllebæk catchment were characterised at 12 localities by means of geological crosssections 10–100 m long, from the margin of the riparian zone adjacent to the cultivated field to the margin of the riparian zone on the opposite side of the stream channel (fig. 3). in each cross-section 8–10 boreholes were hand drilled to depths of 2–4 m and the sediment described at 15 cm intervals. from this geological profiling of the riparian zones, the hydrogeological contact between the stream system and the underlying aquifer was conceptualised according to the typology of groundwater – surface water interaction (gsi; dahl et al. 2007), see fig. 4a. the hydrogeological settings adjacent to the riparian area aquifer were classified at the twelve cross-sections (fig. 4b). combining the regional conceptual model (fig. 2) and the cross-sections, a preliminary delineation of local or regional sandy aquifers 0 20 40 60 80 100 m 4 5 6 7 ox red ox red ew 19 20 0 0 0 0 0 0.1 cultivated field riparian zoneriparian zone cultivated field stream altitude (m) organic-enriched topsoil postglacial freshwater clay postglacial freshwater sand meltwater sand and gravel flow path tile drain redox transition sampling point water table fig. 3. example of a cross section along hagens møllebæk (10 times vertical exaggeration). the profile represents a riparian hydrological type 6 of fig. 4a. nitrate concentrations in no3-n mg/l shown in groundwater, stream water and drain water. the water samples were collected in january 2017; the nitrate concentration in the submerged drain pipe outlet was about 15 mg/l, and 5–6 mg/l in the stream water. 3 1 3/5 5 6/7 6/7 3/5 5 3 1 1/2 3/6 6 • intensive • extensive stream hagens møllebæk catchment monitoring station 3 lateral local / 4 regional 7 unconfined local / 8 regional 5 bottom local / 6 regional 1 disconnected 2 confined riparian hydrogeological type 2 km holocene sand, clay and organic soils clayey till moraine sand and gravel meltwater sand and gravel holocene sand and clay beach dunes 3 a b •••••• • • • • • • • • • fig. 4. riparian zones along hagens møllebæk and its tributaries, divided into riparian hydrogeological types. a: eight conceptual hydrogeological models of the riparian zone with different combinations of deposits with high (white-dotted) and low (grey) permeability (from dahl et al. 2007). b: classification of 12 cross-sections in the study area into types of riparian hydrogeological contacts (see also main text). 19 was obtained (fig. 5). further detailed field instigations are still needed at six of the locations to distinguish between local and regional aquifer characteristics. it is evident that stream reaches along the headwater of the stream system have contact to local shallow aquifers, and that the more central and down-gradient parts of the stream system have contact to a regional sand aquifer. observed exchange between stream water and groundwater stream discharges were measured on the same day in august 2016 at all stations using an ott acoustic digital current meter (ott hydromet gmbh). the groundwater flux to the stream system, quantified as the so-called specific baseflow, was calculated as the change in flow (q in l/sec) between an upstream and downstream location, divided by their distance along the stream. the specific baseflow for the summer day in august 2016 is shown in fig. 5b. the first-order streams provide the lowest inflow rates of groundwater to the stream (0–5 l/sec/km) because of limited or no interaction between the stream system and groundwater aquifers. higher-order (more down-gradient) stream reaches provide 5–25 l/sec/km. along a shorter reach between stations 4 and 5 the tributary loses water through the stream bed to the underlying shallow aquifer. at the twelve intensively monitored stations (fig. 1a), 8–10 piezometers were installed along the cross-sections using metal or polyethylene (peh) pipes with a screen length of 10 cm (metal) and 30 cm (peh). the metal piezometers were pushed into the subsurface with a pneumatic hammer. the peh pipes were manually pushed into the open boreholes for geological characterisation. all piezometers were levelled using the trimble® r8 gps system (vertical accuracy ±16 mm). vertical changes in the hydraulic heads between the water table in the stream (i.e. the water stage) and the intakes of the piezometers about 0.5 m below the stream beds were measured with an accuracy of ±1 cm. negative, vertical, hydraulic gradients indicate recharge to the groundwater aquifer, whereas positive gradients indicate discharge of groundwater into the stream. this made it possible to identify stream reaches where the exchange and direction between the stream and the aquifer can be assessed. the measurements carried out from august 2016 to january 2017 showed similar spatial distributions. this means that the vertical hydraulic gradients were positive along most stream reaches indicating groundwater discharge to the stream during the period august 2016 – january 2017. based on the detailed conceptual understanding of the local hydrogeological setting, the quantitative determination of baseflow and hydraulic gradients, the groundwater–stream interaction at the monitoring stations can be assessed. no or limited groundwater–stream interactions were found at stations 2–4, 6 and 15, while some or significant groundwater–stream interaction is expected to occur at stations 1 and 8–14 between the stream system and an •••••• • • • • • • • • • 2 km 2 km 0–5 10–25 < 0 5–10 hagens m øllebæk •••••• • • • • • • • • • 9 8 7 6 5 4 2 1 14 13 12 10 3 15 11 •••••• • • • • • • • • • 9 8 7 6 5 4 2 1 14 13 12 10 3 15 11 2 km • intensive • extensive local regional 10–25 5–10 0–5 < 0 l/sec/km stream with monitoring station hagens møllebæk catchment monitoring station groundwater aquifers local and regional aquiferslocal and regional aquifers august 2016 october 2016 january 2017 25 mg/l 0 nitrate in stream water (mg/l)specific baseflow in august 2016specific baseflow in august 2016 nitrate in stream water (mg/l) a b c •13 fig. 5. aquifers and nitrate in the hagens møllebæk catchment. a: delineation of local and regional groundwater aquifers underneath the river system in contact with the river bed at the monitoring stations. b: estimated specific base flow for individual stream reaches in august 2016. c: nitrate concentrations in stream water sampled in august and october 2016 and in january 2017. 2020 underlying sandy aquifer of local or regional extent. further investigations of the hydraulic conditions are required at stations 5 and 7 before these can be classified. exchange of nitrate between groundwater, drains and stream water the nitrate transport pathways in riparian zones were assessed by collecting water samples from groundwater in the riparian zones, in drain-pipe outlets, in the stream bed and in stream water. groundwater samples were collected from the piezometers using 100 ml syringes. drain-pipe outlets to the stream were only observed near five of the crosssections (stations 4, 5, 8, 10 and 12). water from drains having outlets above the water stage in the stream was sampled by filling a bottle directly from the outlet. sampling of drain water from outlets below the water stage was more critical. here a 2–3 m long, 5 mm tube was pushed into the drain pipe and a sample collected with a 100 ml syringe. the stream water itself was sampled by filling a bottle of water flowing past the monitoring station. all water samples were analysed for nitrate (no3-n) few hours after collection using a portable photoflex std photometer (wtw gmbh, weilheim). in the cross-section of the riparian zone shown in fig. 3, groundwater samples with nearly 20 mg/l were collected in the oxic zone 25 m east of the stream, decreasing to 0 mg/l underneath the stream bed. normally, nitrate is not degraded in the oxic zone and the observed disappearance of nitrate in the riparian zone is likely due to the presence of microniches in the meltwater sand and gravel layer enriched with organic matter or pyrite, where nitrate degradation may occur. moreover, it is possible that simultaneous discharge of nitrate-free groundwater into the riparian zone and stream bed may dilute the nitrate concentration in the groundwater in the riparian zone. at present we do not possess field data on a sufficiently detailed scale to support this hypothesis. the stream water at the 15 measuring stations contained less than 5–6 mg/l of nitrate in the three sampling rounds in august 2016, october 2016 and january 2017, as expected with the highest values in january 2017 (fig. 5c). at all the stations we have monitored, the riparian zone seems to be efficient in removing nitrate from the groundwater. however, leaching of nitrate from the nearby agricultural areas through drainage pipes can bypass the riparian zone and enter the stream system. a river gauging station was established by the ministry of environment and food of denmark in december 2016 at the same position as station 1 in the present study. the gauging station measures the total river flow and nitrate runoff from the hagens møllebæk catchment to skive fjord. the preliminary results of this study indicate that it is useful to establish a dense monitoring system along streams to complement the new gauging monitoring station. we expect that this type of instrumentation can deliver the necessary insight to locate where nitrate-mitigation measures will be most effective. conclusions after the first six-months’ monitoring period (2016–2017), it seems likely that the nitrate loads to the streams in both summer and winter conditions almost entirely arise from nitrate transported by agricultural drains to hagens møllebæk. however, a longer monitoring period of two to three years is required to conclude this with more confidence. in the hagens møllebæk catchment, riparian zones along the stream reaches with widths from a few metres to 10 m seem to be very efficient in removing nitrate from groundwater before it discharges into the stream. however, drainage pipes allow nitrate-rich water to enter directly into the stream system bypassing the riparian zone. acknowledgements this work was funded by the innovation fund denmark through the project future cropping. the authors wish to thank bo v. iversen and peter k. engesgaard for constructive comments. references dahl, m., nilsson, b., langhoff, j. & refsgaard, j.c. 2007: review of classification systems and new multi-scale typology of groundwater – surface water interaction. journal of hydrology 344, 1–16. hinsby, k., markager, s., kronvang, b., windolf, j., sonnenborg, t.o. & thorling, l. 2012: threshold values and management options for nutrients in a catchment of a temperate estuary with poor ecological status. hydrology and earth system sciences 16, 2663–2683. winter, t.c. 1999: relation of streams, lakes, and wetlands to groundwater flow systems. hydrogeology journal 7, 28–45. authors’ address geological survey of denmark and greenland, øster voldgade 10, copenhagen, denmark. e-mail: bn@geus.dk. geological survey of denmark and greenland bulletin 31, 2014, 35-38 35 palynological and microfossil biostratigraphy and palaeoecology over the paleocene–eocene transition, femern bælt, northern germany patrick alexander richardt and emma sheldon a palynological and micropalaeontological biostratigraphic and palaeoecological investigation has been carried out on the paleocene–eocene transition of core 10.a.057 from the femern bælt (fig. 1). initial investigations of boreholes from the femern bælt indicated that core 10.a.057 included a thick succession of upper paleocene – lower eocene clay (sheldon & nøhr hansen 2010; rambøll arup jv 2011; fig 2). complete paleocene–eocene sections have been described from jylland and the storebælt (heilmann-clausen 1985, laursen & andersen 1997; laursen & king 2000, nielsen et al. 1986), but no detailed studies have been published on these successions from the femern bælt. boreholes were drilled on lolland in denmark, under the femern bælt and on fehmarn island, germany from 2009 to 2011 as part of geological and geophysical investigations performed in preparation for the construction of a fixed road and rail link connecting denmark and germany. the boreholes penetrated campanian – upper eocene strata, overlain by quaternary deposits. borehole 10.a.057 is located in the southern part of the femern bælt (figs 1, 2). geology and palaeoecology the femern bælt is located in the german basin south of the ringkøbing–fyn high. during the late cretaceous, the danish area was characterised by an epicontinental sea resulting in chalk deposition. the sea became more restricted in the early–middle paleocene and several highs bordered the marine area (clausen & huuse 2002). transgression during the selandian resulted in clastic marine sedimentation and the north sea, denmark and the german basin formed a partially enclosed shelf area. during the late paleocene and early eocene, sediment deposition occurred in a relatively deep marine basin, at some distance from the shore. intense volcanic activity caused by the opening of the north atlantic resulted in deposition of ash and tuff layers during this period. the present distribution of the upper part of the palaeogene sediments is a result of erosion and glaciotectonic deformation during the quaternary (fig. 2). in the 10.a.057 core, the very fine-grained clays of the upper paleocene holmehus formation and østerrende clay (informal lithostratigraphic unit of nielsen et al. 1986) are overlain by the lower eocene ølst formation, which is characterised by dark grey clay with abundant layers of black volcanic ash (heilmann-clausen et al. 1985). © 2014 geus. geological survey of denmark and greenland bulletin 31, 35–38. open access: www.geus.dk/publications/bull ringkøbing–fyn high 100 km 14°e 57°n 55° sweden denmark germany german basin norway lolland fehmarn femern bælt 10°e 10.a.057 viborg-1 bovlstrup jylland storebælt north sea stolle klint fig. 1. map of denmark and northern germany showing the location of the planned fixed road and rail link across the femern bælt and the location of borehole 10.a.057 at 54°31.8´n, 11°15.9´e. fehmarn lolland10.a.057 3 km 0 m 50 100 150 quaternary deposits folded and faulted palaeogene clay palaeogene cretaceous fig. 2. sketch south–north cross-section of the femern bælt area showing the location of borehole 10.a.057. 3636 biostratigraphy a total of 19 samples were analysed for dinocysts, diatoms, foraminifers and radiolarians (figs 3, 4). the danish paleocene–eocene zonation of the viborg-1 cored borehole (heilmann-clausen 1985) and the north sea zonation of mudge & bujak (1996) were used for the dinoflagellate cyst stratigraphy. the north sea cenozoic zonation of king (1989) was used for microfossils. dinocysts the samples from 100.49 to 90.31 m are assigned to the areoligera gippingensis acme subzone p5a (mudge & bujak 1996; fig. 3). this subzone represents the uppermost part of zone v4 (heilmann-clausen 1985) and is equivalent to the uppermost part of the holmehus formation (mudge & bujak 1996). p5a is characterised by an acme of a. gippingensis and the presence of eisenackia margarita. the top of p5a is defined by the top of the a. gippingensis acme. nielsen et al. (1986) found low abundances of deflandrea oebisfeldensis in v4 in a borehole from the storebælt but heilmann-clausen (1985) did not find this species in v4. the boundary between v4 and the overlying v5 is tentatively placed, since relatively high abundances of a. gippingensis continue into v5. the samples from 90.31 to 79.42 m are assigned to the e. margarita subzone p5b (v5) based on the last occurrence (lo) of e. margarita (mudge & bujak 1996). e. margarita has its lo in the østerrende clay (nielsen et al. 1986). zone v6, which is characterised by a dominance of the warm-water genus apectodinium and an acme of apectofig. 3. range chart showing the distribution (number of specimens) of dinocysts, diatoms, foraminifers and radiolarians from borehole 10.a.057. p: pyrite, a: ash layer. observed outside counting very rare (1) rare (2–4) common (5–14) abundant (15–49) dominant (>50) c hr on os tr at igr ap hy ea rly e oc en e la te p ale oc en e li th os tr at igr ap hy ø lst f or m at io n h ol m eh us f or m at io n ø st er re nd e c lay d an ish v ib or g di no cy st z on at io n v7 e1 b v5 v4 p5 a p5 b n or th s ea m icr of os sil z on es m ud ge & b uj ak 1 99 4, bu jak & m ud ge 1 99 6 n sp 4 / n sa 2 n sp 3 / n sa 1b ac ho m os ph ae ra sp p. co rd os ph ae rid ium sp p. de fla nd re a oe bis fe lde ns is di no cy st sp p. m icr od ini um c f. or na tu m pa lae ot et ra din ium m inu sc ulu m sp ini fe rit es ra m os us sp ini fe rit es sp p. h ys tri ch os ph ae rid ium tu bif er um im pa gid ini um sp p. ol igo sp ha er idi um co m ple x th ala ss iph or a de lic at a ac ho m os ph ae ra cr as sip ell is gl ap hy ro cy sta o rd ina ta gl ap hy ro cy sta sp p. gl ap hy ro cy sta d iva ric at a h ys tri ch os ph ae rid ium tu bif er um b re vis pin um al iso cy sta sp . 2 cr ibr op er idi niu m sp p. op er cu lod ini um sp p. cr ibr op er idi niu m te nu ita bu lat um u ni nd en tif ie d pe rid in oi d cy st s ac ho m os ph ae ra a lci co rn u ar eo lig er a gip pin ge ns is ac hil leo din ium la tis pin os um ei se na ck ia m ar ga rit a pa lae oc ys to din ium lid iae ar eo lig er a sp p. dinoflagellate cysts di at om sp p. fe ne str ell a an tiq ua co sc ino dis cu s m or sia nu s m oe lle ri th ala ss ior iro ps is wi tti an a au lac od isc us a llo rg ei di at om sp p. fl at si lic eo us h em iau lus sp p. tr ina cr ia re gin a di at om sp p. fl at p yr itis ed fe ne str ell a an tiq ua (s m all ) diatoms sp iro ple cta m m ina sp ec ta bil is sp iro ple cta m m ina sp p. ag glu t. f or am ini fe ra in de t am m od isc us cr et ac eu s gl om os pir a ch ar oid es rh ab da m m ina sp p. ba th ys iph on m icr or ha ph idu s ba th ys iph on sp p. cy cla m m ina a m ple cte ns cy cla m m ina ro tu nd ido rs at a cy sta m m ina p au cil oc ula ta h ap lop hr ag m oid es sp p. h ap lop hr ag m oid es w alt er i h or m os ina sp p. la br os pir a sc itu la m ar ss on ell a ox yc on a re cu rv oid es sp p. re op ha x sp p. rh ab da m m ina ro bu sta agglutinating benthic foraminifers ra di ol ar ian s ce no dis cu s s pp . 100 95 90 85 80 75 70 65 60 55 p p p p p p p p p p a a a p d ep th (m ) 37 dinium augustum, was not found in this study. a missing core section from 79.42 to 73.36 m could represent zone v6, but v6 has not been observed previously in the femern bælt area (c. heilmann-clausen & h. nøhr-hansen, personal communication 2013). however, v6 is present in the stolle klint clay, north-western jylland. the base of the apectodinium acme in the earliest eocene is a global event and is coupled with a carbon isotope excursion denoting the start of the paleocene–eocene thermal maximum (petm). the samples from 73.36 to 51.81 m are referred to early eocene d. oebisfeldensis acme subzone e1b (upper v7; bujak & mudge 1994). the interval from 66.18 to 51.81 m is characterised by a minor acme of glaphyrocysta divaricata and a high abundance of microdinium cf. ornatum (fig. 3). zone v7 is found in the upper part of the ølst formation (heilmann-clausen 1985). foraminifers and diatoms the assemblages mainly consist of poorly preserved agglutinating benthic foraminifers and diatoms. samples from 100.49 to 79.42 m are assigned to zone nsa1b. most samples only contain few foraminifers, but two samples from the upper c. 2 m contain rich faunas with spiroplectammina spectabilis, labrospira scitula, ammodiscus cretaceus, glomospira charoides, marsonella oxycona, cystammina pauciloculata, recurvoides spp., hormosina spp., rhabdammina robusta, cyclammina rotundidorsata, haplophragmoides walteri, bathysiphon spp. and cyclammina amplectens. this assemblage is known as the ‘rhabdammina biofacies’. the shift from low to higher diversity benthic assemblages in nsa1b was also noted in the bovlstrup borehole, eastern jylland (laursen & andersen 1997) and was interpreted as a shift from very poor to slightly improved life conditions on the sea floor. a low-diversity diatom flora with pyritised fenestrella antiqua (var. small) and diatom spp. (flat) is also present, in addition to sponge debris and radiolarians (cenodiscus spp.). nsa1 is assigned to the holmehus formation in denmark (king 1989). samples from 73.36 to 51.81 m are assigned to zone nsp4. the assemblages are dominated by resting spores of centric diatoms, comprising coscinodiscus morsianus moelleri, fenestrella antiqua, diatom spp. (flat), thalassiophora wittiana, trinacria regina, aulacodiscus allorgei and hemiaulus spp. from 66.18 to 56.87 m pyritised and translucent diatoms occur in equal numbers; above and below this level only pyritised specimens occur. this difference in preservation is probably due to variations in the oxygen level in the water column and the amount of sulphide present in the sediment (de jonghe et al. 2011). zone nsp4 also includes bryozoan fig. 4. selected fossils: a–c and f–i: foraminifers, d, e, j, k: diatoms, l–r: dinocysts. a: ammodiscus cretaceous. b: cyclammina amplectens. c: cystammina pauciloculata. d: coscinodiscus morsianus moelleri. e: fenestrella antiqua. f: glomospira charoides. g: haplophragmoides walteri. h: labrospira scitula (front). i: labrospira scitula (side). j: trinacria regina. k: trinacria regina (siliceous). l: microdinium cf. ornatum. m: piece of dinocyst. n: areoligera gippingensis. o: deflandrea oebisfeldensis. p: eisenackia margarita. q: hystrichospharidium tubiferum. r: unidentifiable peridinoid cyst. scale bars: 100 µm (a–k), 20 µm (l–r). a b c d e f g h i j k m n o p l q r 3838 fragments, fish teeth, inoceramus fragments, sponge spicules and rare agglutinating foraminifers. nsp4 is assigned to the early eocene ølst formation (king 1989). discussion and conclusions a dominance of agglutinating foraminifers of the ‘rhabdammina biofacies’ in subzone nsa1b suggests a middle to lower bathyal palaeoenvironment characterised by restricted water circulation, with low oxygen levels and a reducing environment at the sea floor (jones & charnock 1985; king 1989). the low oxygen level may have led to decreasing degradation of organic matter by bacteria and benthic organisms, giving rise to the dark grey colour of the holmehus formation compared with those described by heilmann-clausen et al. (1985). a relatively high abundance of the supposed heterotrophic dinoflagellate deflandrea oebisfeldensis and the rich diatom flora in the ølst formation may be due to increased productivity in the surface layers perhaps due to enhanced upwelling in coastal areas. the rich diatom flora, preserved as resting spores, also suggests stressed sea-surface and seabed conditions, perhaps due to volcanic ash falls and periods of anoxia (bidgood et al. 1999). the scarcity of agglutinating foraminifers is probably due to reducing conditions at the sea floor. schiøler et al. (2007) suggested that the balder formation (the north sea equivalent of the ølst formation) was deposited in a restricted marine environment at upper bathyal depths with dysoxic to anoxic bottom conditions. in core 10.a.057, palynological biostratigraphy indicates that zone v6 is absent, either due to erosion or non-deposition, therefore the paleocene–eocene thermal maximum is not preserved at this location. a hiatus with v6 missing was noted in the lillebælt area (heilmann-clausen et al. 1985), supporting the femern bælt data. however it is possible that zone v6 is found in the missing core interval between 79.42 and 73.36 m. references bidgood, m.d., mitlehner, a.g., jones, g.d. & jutson, d.j. 1999: towards a stable and agreed nomenclature for north sea tertiary diatom floras – the ‘coscinodiscus’ problem. in: jones, r.w. & simmons, m.d. (eds): biostratigraphy in production and development geology. geological society (london) special publications 152, 139–153. bujak, j. & mudge, d. 1994: a high-resolution north sea eocene dinocyst zonation. journal of the geological society 151, 449-462. clausen, o.r. & huuse, m. 2002: mid-paleocene palaeogeograhpy of the danish area. bulletin of the geological society of denmark 49, 171–186. de jonghe, a., hart, m.b., grimes, s.t., mitlehner, a.g., price, g.d. & smart c.w. 2011: middle eocene diatoms from whitecliff bay, isle of wight, england: stratigraphy and preservation. proceedings of the geologists’ association 122, 472–483. heilmann-clausen, c. 1985: dinoflagellate stratigraphy of the uppermost danian to ypresian in the viborg 1 borehole, central jylland, denmark. danmarks geologiske undersøgelse serie a 7, 69 pp. heilmann-clausen, c., nielsen, o.b. & gersner, f. 1985: lithostratigraphy and depositional environments in the upper paleocene and eocene of denmark. bulletin of the geological society of denmark 33, 287–323. jones r.w. & charnock, m.a. 1985: ‘morphogroups’ of agglutinating foraminifera: their life positions, feeding habits and potential applicability in (paleo)ecological studies. revue de paléobiologie 4, 311–320. king, c. 1989. cenozoic of the north sea. in: jenkins, d.g. & murray, j.w. (eds): stratigraphical atlas of fossil foraminifera, 418–489. chichester: ellis horwood. laursen, g.v. & andersen, s.b. 1997: a late palaeocene–early eocene benthic foraminiferal record from bovlstrup, denmark, showing a remarkable agglutinated fauna. journal of micropalaeontology 16, 19–29. laursen, g.v. & king, c. 2000: preliminary results of a foraminiferal analysis of a core from østerrende, denmark. geologiska föreningen i stockholm förhandlingar (gff) 122, 92 only. mudge, d.c. & bujak, j.p. 1996: palaeocene biostratigraphy and sequence stratigraphy of the uk central north sea. marine and petroleum geology 13, 295–312. nielsen, o.b., baumann, j., zhang, d., heilmann-clausen, c. & larsen, g. 1986: tertiary deposits in store bælt. the tertiary section of borehole d.g.i. 83101, østerenden, storebælt, denmark. in: møller, j.t. (ed.): twentyfive years of geology in aarhus. geoskrifter 24, 237-249. department of geoscience, aarhus university, denmark. rambøll arup jv 2011: summary of geological conditions. geotechnical data report 01.3-002, 53 pp. virum: femern a/s. schiøler, p. et al. 2007: lithostratigraphy of the palaeogene – lower neogene succession of the danish north sea. geological survey of denmark and greenland bulletin 12, 77 pp. sheldon, e. & nøhr-hansen, h. 2010: fehmarn belt fixed link pre-quaternary biostratigraphy – a final status report for ramboll arup joint venture. danmarks og grønlands geologiske undersøgelse rapport 2010/134, 53 pp. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: par@geus.dk geological survey of denmark and greenland bulletin 1, 231-246 danish central graben north sea baltic sea 10°e8°e 12°e 14°e 16°e 58°n 57°n 56°n 55°n 6°e4°e sweden the netherlands germany norway denmark uk fault national boundaries bornholm 50 km central graben 233 mapping of the danish central graben has provided a coherent homogeneous database for evaluating the main features of structural evolution during the late jurassic and the early cretaceous (britze et al. 1995a, b, c, d). this mapping has highlighted the necessity for the definition of certain new structural elements and clarification of the established terminology. the aim of this article therefore is to present a brief description of the important tectonic elements of the area in the late jurassic – early cretaceous, illustrated by reduced versions of maps published at a scale of 1:200 000 (britze et al. 1995a, b, c, d). in particular, we distinguish between the elements that were active during the late jurassic – earliest cretaceous and those that came into existence during the deposition of the cromer knoll group. this then provides the framework for subsequent detailed articles on the upper jurassic of the danish central graben (andsbjerg & dybkjær 2003, this volume; johannessen 2003, this volume; møller & rasmussen 2003, this volume). the introduction to the upper jurassic – lower cretaceous structural geology of the danish central graben is preceded by a brief account of the mapped horizons (fig. 1) and the mapping procedures adopted. as discussed further below, the seismic boundary between rocks largely of jurassic age and the overlying lower cretaceous cromer knoll group is located upper jurassic – lower cretaceous of the danish central graben: structural framework and nomenclature peter japsen, peter britze and claus andersen the danish central graben is part of the mainly late jurassic complex of grabens in the central and southern north sea which form the central graben. the tectonic elements of the danish central graben in the late jurassic are outlined and compared to those in the early cretaceous based on reduced versions of published maps (1:200 000), compiled on the basis of all 1994 public domain seismic and well data. the tail end graben, a half-graben which stretches for about 90 km along the east north sea high, is the dominant late jurassic structural feature. the rosa basin (new name) is a narrow, north–south-trending basin extending from the south-western part of the tail end graben. the tail end graben ceased to exist as a coherent structural element during the early cretaceous and developed into three separate depocentres: the iris and gulnare basins to the north and the roar basin to the south (new names). the early cretaceous saw a shift from subsidence focused along the east north sea high during the late jurassic to a more even distribution of minor basins within the danish central graben. the depth to the top of the upper jurassic – lowermost cretaceous farsund formation reaches a maximum of 4800 m in the northern part of the study area, while the depth to the base of the upper jurassic reaches 7500 m in the tail end graben, where the upper jurassic attains a maximum thickness of 3600 m. the lower cretaceous cromer knoll group attains a maximum thickness of 1100 m in the outer rough basin. keywords: north sea, danish central graben, upper jurassic, lower cretaceous, isochore maps, structure contour maps, structural nomenclature geological survey of denmark and greenland, geocenter copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: pj@geus.dk geological survey of denmark and greenland bulletin 1, 233–246 (2003) © geus, 2003 at the top of the farsund formation which extends up into the lowermost cretaceous (the boundary is of late ryazanian age). thus, the term ‘upper/late jurassic’ is used broadly in this paper so as to include the lowermost/earliest cretaceous which genetically forms part of the upper jurassic succession. regional setting the danish central graben consists of a system of generally nnw–sse-trending half-grabens bounded by the coffee soil fault to the east towards the footwall block of the east north sea block, and by the mid north sea high to the west (figs 2–6). the rifting, which involved high rates of crustal stretching (vejbæk 1992; ineson et al. 2003, this volume), commenced during the bajocian (johannesen & andsbjerg 1993). the syn-rift sedimentary fill is mudstone-dominated and is very rich in organic matter at certain levels (damtoft et al. 1992). the early development was characterised by fault-controlled subsidence and deposition in the eastern part, especially along north–south-trending segments of the coffee soil fault (korstgård et al. 1993). during the kimmeridgian, the dominant tectonic trend shifted from the north–south fault direction inherited from the pre-jurassic to a dominant nw–se direction (møller & rasmussen 2003, this volume); the depocentres shifted westwards and deposition gradually covered larger areas. despite the overall extensional tectonic regime, compressional features have been reported, caused by oblique-slip movements between different graben segments or by re-adjustments at boundaries between opposite-dipping fault blocks (sundsbø & megson 1993; rasmussen 1995). the extensional tectonic regime of the late jurassic continued in the early cretaceous, albeit with much reduced subsidence rates. although subsidence patterns to some extent followed those of the late jurassic, there was a shift with new depocentres developing to the west (figs 7, 8). the lower cretaceous deposits accumulated in local depocentres, partly as a result of reduced sediment supply, and partly due to accentuation of ridges and structural highs separating the subbasins. block faulting gradually ceased after the hauterivian, giving way to regional subsidence and mild inversion movements (vejbæk 1986). central graben versus central trough the term ‘central graben’ has gained wide acceptance (ziegler 1990), in preference to ‘central trough’, a term that was first proposed by rønnevik et al. (1975). their argument for adopting the term central trough for the “graben system of permian–cretaceous age” was that “the tectonic axis of the graben system does not coincide with the deepest part of the depocentre and this is the reason why the name trough is preferred over graben” (rønnevik et al. 1975, p. 5). in contrast, gowers & sæbøe (1985, p. 313) stated that “the central trough should be treated as a broad nw-trending zone of subsidence with some local highs only when discussing post midcretaceous geology”. the central trough is consequently identified on a structural map of the base of the cenozoic by gowers & sæbøe (1985, fig. 1). when considering the 234 ‘top chalk’ ‘base chalk’ ‘base cretaceous’ post chalk group chalk group cromer knoll group upper jurassic units pre-jurassic units l. and m. jurassic units britze et al. 1995a, b, c, dlithostratigraphy chronostratigraphy top jurassic (fig. 3) base upper jur. (fig. 4) base upper jur. p. c re ta ce ou s ju ra ss ic c en oz . m es oz oi c n .q . l. –m . u pp er lo w er u p. fig. 1. generalised stratigraphic scheme indicating the relative positions of the horizons mapped by britze et al. (1995a, b, c, d) and the surfaces shown in figures 3 and 4. whereas the surfaces mapped by britze et al. (1995a, b, c, d) are defined for every point in the study area, those of figures 3 and 4 are only defined where the upper jurassic is defined in the central graben. inverted commas (e.g. ‘top chalk’) are used to indicate that the referenced surface may not be present throughout (i.e. the chalk may locally be absent). the main lithostratigraphic units applied in the mapping are the post chalk group (cenozoic excluding the danian; nielsen & japsen 1991), the chalk group (upper cretaceous – danian; lieberkind et al. 1982), the cromer knoll group (deegan & scull 1977) and the term ‘upper jurassic units’ comprising the lola, heno, farsund and poul formations. for details of the lithostratigraphic subdivision, see michelsen et al. (2003, this volume). jur, jurassic; l, lower; m, middle; up, upper; cenoz, cenozoic; p, palaeogene; n, neogene; q, quaternary. 235 pre mid-cretaceous geology, the recommendation of gowers & sæbøe (1985) was to refer to the individual structural elements within the area of the central trough. we find the use of the term central trough confusing, and propose that the term central graben is used for the complex of grabens and half-grabens of mainly late jurassic age in the central and southern north sea. furthermore, we recommend that the term north sea basin (ziegler 1990) is adopted when referring to the regional subsidence pattern of late cretaceous and cenozoic post-rift sedimentation that was aligned over the former mesozoic grabens. subsurface mapping of the danish central graben the structural depth and isochore maps presented here are reduced and simplified versions of a suite of fifteen maps published at a scale of 1:200 000 and covering up to 13 000 km2 in the western part of the danish north sea sector (britze et al. 1995a, b, c, d). the maps are part of the outcome of an integrated mapping project with the aim of modelling and mapping interval velocities and depths of the main post-middle jurassic units in the danish central graben. the study is based on 4°e 56°n 55°30'n 5°e e. rosa-3 elin-1 nora-1 liva-1 eg-1 iris-1 tordenskjold-1 jeppe-1 amalie-1 lulu-1 gulnare-1 cleo-1 l-1 mona-1 karl-1 w. lulu-3 john ravn-2 ravn-1 falk-1 elly-1 elly-2 b-1 w-1 bo-1 boje-1 north jens-1 edna-1 e-1 jens-1 adda-1 deep adda-1 v-1 ugle-1 g-1 s. e. igor-1 emma-1m-8 m-1 a-2 anne-3 o-1 alma-1 vagn tove nils john flank-1 john u-1 deep gorm-1 ruth m. rosa e. rosa e. rosa flank-1 sten-1 kim-1 lone-1 gert-1 4 2 3 gwen-2 q-1 t-3 diamant-1 i-1 1 2 2 4 n. arne well location map 1:880 000 utm zone 31 normal fault reverse fault fault plane name of salt diapir jurassic thin or absent well reaching top jurassic john 25 km uk n dk g nl fig. 2. wells reaching top jurassic. well database for the mapping incorporating all 1994 released wells plus the alma-1 and amalie-1 wells. national sectors of the north sea: dk, denmark; g, germany; n, norway; nl, the netherlands; uk, united kingdom. 236 well data and time structure maps of well-defined marker horizons illustrating both the syn-rift and post-rift development of the area. previous work the central graben was recognised as one of the main geological provinces in denmark by rasmussen (1978) based on seismic mapping by j.c. baartman. the first structural outline of the danish central graben was presented by andersen et al. (1982), and the structural evolution of the danish and norwegian parts of the central graben was later discussed by gowers & sæbøe (1985). detailed analysis of the structural development and definition of the structural nomenclature for the danish central graben was presented by møller (1986) for the middle and upper jurassic and by vejbæk (1986) for the lower cretaceous. the structure of the central 5°e 56°n 55°30'n 4°e normal fault at ‘base upper jurassic’ level reverse fault at ‘base upper jurassic’ level fault plane jurassic thin or absent well reaching base upper jurassic base upper jurassic depth structure map contour interval 400 m 1:880 000 utm zone 31 depth in metres below mean sea level 2000 4000 6000 25 km fig. 3. base upper jurassic structural depth map showing depth to the base of the upper jurassic units (fig. 1). modified after britze et al. (1995d). 237 graben in the norwegian and northernmost danish sectors was further discussed by gowers et al. (1993). four two-way time maps of the jurassic in the danish central graben were published by møller (1986): near base middle jurassic/top pre-jurassic, oxfordian– kimmeridgian boundary, late kimmeridgian marker and near top jurassic/base cretaceous (1:500 000) (mapped surfaces are identified by italics in this paper). in addition, a map (1:800 000) was presented showing the thickness in metres of the middle and upper jurassic. database the mapped area covers the danish central graben and part of the eastern footwall block, the east north sea block, which forms part of the ringkøbing–fyn high system of elevated basement blocks (figs 5, 7). all 1994 public domain petroleum industry seismic and well data acquired on danish territory form the basis for the study, with the addition of data from the alma-1 and amalie-1 wells. the well database comprises 96 released 4°e 5°e 56°n 55°30'n depth in metres below mean sea level top jurassic depth structure map contour interval 200 m 1:880 000 utm zone 31 normal fault at ‘base cretaceous’ level reverse fault at ‘base cretaceous’ level fault plane jurassic thin or absent well reaching top jurassic depth in metres below mean sea level 2000 3000 4000 5000 25 km fig. 4. top jurassic structural depth map showing depth to the top of the upper jurassic units (fig. 1). modified after britze et al. (1995c). 238 exploration and appraisal wells drilled, as a minimum, into the chalk group (upper cretaceous – danian, lieberkind et al. 1982). of these wells, 62 drilled into upper jurassic units (figs 1, 2). the lithostratigraphic subdivision of most of the wells is presented in nielsen & japsen (1991). the seismic database varies in quality from 1979 2d sections to 1988 3d data (rasmussen 1995; kristensen et al. 1995), and comprises regional speculative surveys acquired during the early and mid1980s and proprietary surveys acquired by mærsk olie og gas as. seismic markers in the eastern part of the central graben, the upper jurassic graben fill conformably overlies the middle jurassic. here the base upper jurassic seismic sequence boundary is picked in a trough above a strong peak that usually shows significant lateral continuity. this seismic boundary marks a general increase in acoustic impedance. the trough–peak relationship is believed to be enhanced by the interference of alternating sandstones, shales and coal beds in the uppermost part of 5°e 56°n 55°30'n 4°e mid north sea high feda graben gert ridge gertrud graben heno plateau inge high mandal high piggvar terrace søgne basin m ads high poul plateau salt dome province tail end graben r osa basin arne–elin graben coffee soil fault ringkøbing–fyn high east north sea block inge high late jurassic structural elements 1:880 000 utm zone 31 normal fault at ‘base upper jurassic’ level reverse fault at ‘base upper jurassic’ level fault plane name of late jurassic structural element jurassic thin or absent well reaching top jurassic 25 km skrubbe fault fig. 5. late jurassic structural elements (britze et al. 1995d). based on the upper jurassic isochore (fig. 6) and møller (1986). 239 the middle jurassic. further westwards, where the upper jurassic overlies a pre-jurassic substratum, the seismic marker is recognised as a pronounced unconformity. interpretation of the base upper jurassic surface is complicated due to the great depth of burial (up to 6 seconds two-way time), the structural complexity, and local interference of multiples caused by overlying, thick high-velocity chalk deposits, for example on the heno plateau where lower cretaceous strata are absent. in basinal areas, the top jurassic marker is an easily identified seismic event marking the top of the low-velocity farsund formation of late jurassic – earliest cretaceous age (jensen et al. 1986). towards the basin margins, the overlying sequence onlaps top jurassic. locally the upper jurassic sediments have been eroded and the top jurassic marker is recognised as an unconformity. seismic mapping four horizons were mapped throughout the study area to illustrate the structural evolution and to undertake 5°e 56°n 55°30'n 4°e 25 km upper jurassic isochore map contour interval 400 m 1:880 000 utm zone 31 normal fault at ‘base upper jurassic’ level reverse fault at ‘base upper jurassic’ level fault plane jurassic thin or absent well reaching base upper jurassic thickness in metres 0 2000 4000 fig 6. upper jurassic isochore showing the thickness of the upper jurassic units (fig. 1). modified after britze et al. (1995d). 240 layer-cake depth conversion (fig. 1). the maps are thus not only defined by the presence of a given lithological unit (e.g. top chalk group), but rather by the base of all sediments above the unit (e.g. the base of the cenozoic deposits, excluding the danian). however, a simple nomenclature was chosen for the mapped horizons corresponding to the everyday usage (‘top chalk’) rather than applying a strict terminology, and inverted commas were adopted to indicate that the referenced surface may not be present throughout (i.e. the chalk group may locally be absent). furthermore, a horizon name may represent a stratigraphic simplification; for example, the uppermost farsund formation, bounded by the ‘top jurassic’ surface, may include rocks of earliest cretaceous (ryazanian) age. the four horizons mapped are ‘top chalk’ – base of the cenozoic deposits (excluding the danian), ‘base chalk’ – base of the post-early cretaceous deposits, ‘base cretaceous’ – base of the post-jurassic deposits and ‘base upper jurassic’ – base of the post-middle jurassic deposits (britze et al. 1995a, b, c, d). the ‘base cretaceous’ time map is an extension and revision of the near top jurassic/base cretaceous time map, published by møller (1986). the base upper jurassic time map is an extension and revision of the near base middle jurassic/top pre-jurassic time map (møller 1986). the map was merged with the top middle jurassic structure map of the søgne basin – northern tail end graben by korstgård et al. (1993). both time maps were incorporated with unpublished regional interpretations and detailed mapping of hydrocarbon fields undertaken by the geological survey of denmark and greenland. the drafts of the time structure maps were compiled manually at a scale of 1:100 000. the final maps were produced digitally with the zmap plus mapping system using a 200 m gridding interval. the base upper jurassic and top jurassic depth maps shown in figures 3 and 4 are the depth-converted images of two structural travel-time maps: ‘base cretaceous’ (britze et al. 1995c) and ‘base upper jurassic’ (britze et al. 1995d). however, in order to focus on the jurassic deposits, these maps are only defined where the upper jurassic is present in the central graben, whereas the original maps are defined throughout the area. the upper jurassic isochore map is thus equivalent to the thickness of the upper jurassic units comprising the lola, heno, farsund and poul formations (fig. 1). depth conversion the depths to the surfaces are calculated by multi-layer depth conversion based on a study of different methods by japsen (1994). the depth to top jurassic is calculated by adding the thicknesses of the chalk group and the cromer knoll group (deegan & scull 1977) to the depth to the ‘top chalk’ surface (britze et al. 1995a, b, c). the thickness of the chalk overburden equivalent to the post chalk group (cenozoic excluding danian, nielsen & japsen 1991) is calculated by subdividing this layer into two at the mid-miocene unconformity which represents the top of the overpressured zone (japsen 1994). the thickness of each of these units are calculated by velocity-anomaly depth conversion whereby the thickness of each layer is calculated sequentially from the top downwards from the seismic travel-time thickness by assuming that the velocity of the layer increases linearly with depth (japsen 1993). laterally, however, velocity is calibrated to match well data. the velocity-anomaly map, the seismic traveltime maps, and the linear velocity parameters for each layer constitute input for velocity-anomaly depth conversion. the generated output are depth and intervalvelocity maps. the velocity-depth gradient applied for the upper and lower parts of the post chalk group is 0.4 and 0.2 m/sec/m, respectively, and for the chalk group and cromer knoll group 0.7 and 0.5 m/sec/m, respectively (japsen 1994). the depth to base upper jurassic is calculated by adding the thickness of the upper jurassic to the depth to the top jurassic surface. the depth to base upper jurassic is thus in total based on a five-layer depth conversion. no easily applied relationships between velocity and depth or travel time were found that could be used to predict the interval velocity of the upper jurassic (japsen 1994). this is believed to be due to the complex geological and physical conditions as well as the restricted well database for this interval. the thickness of the upper jurassic is consequently calculated by multiplying the travel time of that unit by the interval velocity of the layer (fig. 6). results upper jurassic sediments are present throughout the main part of the danish central graben except over the inge, mads and mandal highs and a number of salt diapirs (see discussion on the structural nomenclature below): east rosa (the dagmar field), john, middle 241 rosa (the rolf field), nils (the regnar field), north arne (the svend field), ruth (the skjold field), tove and vagn (fig. 2). upper jurassic sediments are thin or absent on the east north sea block. a thin succession of late jurassic age was encountered in the l-1 well, on the south-western flank of the norwegian–danish basin (figs 2, 7); the thickness of this succession is included in the cromer knoll group isochore (fig. 8). the interval-velocity map of the upper jurassic represents contoured data from 54 wells of which 10 encountered more than 400 m of upper jurassic strata without penetrating the entire succession (britze et al. 1995d). the latter wells are included to define the velocity field of the thick upper jurassic sequence in the tail end graben. the interval velocity of the upper jurassic is very low relative to depth in the tail end graben: 2.6 km/sec in the nora-1 well, which drilled through the upper jurassic interval in the depth range from 3500 to 4400 m below sea level (well locations are given on fig. 2). low velocities (< 2.6 km/sec) recorded in the northern part of the tail end graben are believed to be due to extreme overpressure caused by gas generation (japsen 1994). a minimum value of 2.4 km/sec was recorded in the north jens-1 well. velocities are relatively high in wells towards the west of the mapped area and adjacent to parts of the coffee soil fault where deep-water sands originating from the east north sea block are found in the uppermost part of the jurassic (damtoft et al. 1992). the maximum interval velocity was recorded in the ugle-1 well (4.0 km/sec). depth to base upper jurassic ranges from 2100 m in the john flank-1 well to 7500 m in the deepest parts of the tail end graben, where the upper jurassic attains a maximum thickness of 3600 m of which a maximum of 1400 m has been drilled (the g-1 well) (figs 3, 6). depth to top jurassic ranges from 1700 m in the john flank-1 well to 4800 m at the base of several early cretaceous depocentres in the northern parts of the central graben (fig. 4). the maximum thicknesses of the cromer knoll group are found in the outer rough basin (1100 m; figs 7, 8). thicknesses of more than 800 m are estimated in the ål, iris and roar basins and in the feda, gertrud and arne–elin grabens. late jurassic structural elements in the following, a brief presentation of the late jurassic structural elements is given with special emphasis on differences relative to the description by møller (1986). the spatial relationship between these elements (fig. 5) are depicted on the upper jurassic isochore map (fig. 6). the terminology follows that of møller (1986) where reference to other studies is not indicated (table 1). the tail end graben (andersen et al. 1982) is the dominant late jurassic structural element in the danish central graben. it is a half-graben bounded to the east by the generally nw–se-trending segments of the coffee soil fault. the tail end graben accumulated up to 3600 m of sediments during the late jurassic, and it stretches over a length of about 90 km. thicknesses in excess of 2000 m are estimated to be present within an area from the gulnare-1 well in the north to the g-1 well in the south. the depocentre of the tail end graben shifted westwards during the late jurassic while faulting shifted from north–south to nnw–sse trends (møller & rasmussen 2003, this volume). towards the north, the tail end graben grades into the piggvar terrace (gowers et al. 1993), and the north–south oriented danish part of the søgne basin (gowers & sæbøe 1985). in the south, the tail end graben extends into the narål basin gowers et al. (1993) x arne–elin graben møller (1986), vejbæk (1986) x x central graben rasmussen (1978) x x east north sea block rasmussen (1978) x x feda graben gowers & sæbøe (1985) x x gert ridge møller (1986), vejbæk (1986) x x gertrud graben møller (1986), vejbæk (1986) x x gulnare basin new name x heno plateau møller (1986), vejbæk (1986) x x inge high møller (1986), vejbæk (1986) x x iris basin new name x mads high møller (1986), vejbæk (1986) x x mandal high rønnevik et al. (1975) x x mid north sea high rasmussen (1978) x x outer rough basin gowers & sæbøe (1985) x piggvar terrace gowers et al. (1993) x pollerne ridge vejbæk (1986) x poul plateau møller (1986) x x ringkøbing–fyn high rasmussen (1978) x x roar basin new name x rosa basin new name x salt dome province møller (1986), vejbæk (1986) x x skrubbe fault gowers & sæbøe (1985) x x søgne basin gowers & sæbøe (1985) x tail end graben andersen et al. (1982) x e. cret. : early cretaceous l. jur. : late jurassic structural element reference age l. jur. e. cret. table 1. structural nomenclature for the danish central graben row north–south oriented rosa basin (new name) and grades into the salt dome province, a region dominated by halokinetic features. the poul plateau with a reduced thickness of upper jurassic, is adjacent to a shift in the orientation of the coffee soil fault. the piggvar terrace (gowers et al. 1993) is a narrow, nw–se-trending segment separated by major normal faults from the mandal high (rønnevik et al. 1975) to the east and the gertrud graben to the west; it extends further north into norwegian waters. upper jurassic thicknesses reach 800 m on the piggvar terrace adjacent to the mandal high. the søgne basin (gowers & sæbøe 1985) is a halfgraben that has a depocentre along the coffee soil fault in the north-eastern part of the danish central graben and continues into norwegian waters. its evolution is similar to the salt dome province with relatively thick triassic – lower upper jurassic sequences (møller 1986). a hiatus spanning the latest jurassic increases in magnitude towards the mandal high where the jurassic is 242 john 5°e 56°n 55°30'n 4°e norwegian–danish basin outer rough basin ålbasin skrubbe fault inge high feda graben mads high heno plateau roar basin gertrud graben iris basin gulnare basin mandal high arne–elin graben pollerne ridge poul plateau c offee soil fault mid north sea high ringkøbing–fyn high east north sea block salt dome province gert ridge early cretaceous structural elements 1:880 000 utm zone 31 inge high normal fault at ‘base cretaceous’ level reverse fault at ‘base cretaceous’ level fault plane name of early cretaceous structural element cromer knoll group thin or absent well reaching top jurassic 25 km fig. 7. early cretaceous structural elements (britze et al. 1995c). based on the cromer knoll group isochore (fig. 7) and vejbæk (1986). absent (andsbjerg & dybkjær 2003, this volume). however, the preserved pre-kimmeridgian jurassic sequences in the søgne basin show no evidence of any syndepositional movements, an observation that suggests post-kimmeridgian rotation of the mandal high – søgne basin area. the rosa basin (new name) is a narrow, north–southtrending basin that extends from the south-western part of the tail end graben into the salt dome province. the rosa basin is limited by normal faults to the east and to the west, and the thickness of the upper jurassic exceeds 2000 m along the eastern bounding fault. the basin is named after the east rosa salt diapir to the west. the feda graben that continues north into norwegian waters is the most prominent feature in the north-western part of the mapped area; it dips towards the west, where up to 1600 m of upper jurassic sediments accumulated along the skrubbe fault (gowers et al. 1993) towards the mid north sea and inge highs. it is separated from the gertrud graben which dips to the east 243 4°e 5°e 56°n 55°30'n 500 thickness in metres cromer knoll group isochore map contour interval 400 m 1:880 000 utm zone 31 normal fault at ‘base cretaceous’ level reverse fault at ‘base cretaceous’ level fault plane cromer knoll group thin or absent well reaching top jurassic 0 1000 25 km fig. 8. cromer knoll group isochore. note that although the cromer knoll group is not recognised outside the central graben, the map also depicts the thickness of lower cretaceous strata (including a thin upper jurassic section) in the south-western part of the norwegian–danish basin. modified from britze et al. (1995c). by the narrow gert ridge that is interpreted to have come into existence as a high during the earliest cretaceous while the uppermost farsund formation was being deposited (møller & rasmussen 2003, this volume). towards the south, these depocentres grade into the heno plateau, characterised by intermediate thicknesses of upper jurassic sediments, and dominated by a mosaic of minor fault blocks. the arne–elin graben separates the heno plateau from the tail end graben. it is a pull-apart basin, related to left-lateral oblique-slip movements according to vejbæk (1986) and korstgård et al. (1993); the bounding faults were inverted during the cretaceous (møller 1986). the mid north sea high (rasmussen 1978) is covered by a thin veneer of upper jurassic sediments, and constitutes the western part of the danish central graben. it is bounded to the east by the inge and mads highs where the jurassic is absent, and is characterised by a large number of minor, basement-attached faults. early cretaceous structural elements in the following, a brief presentation of the early cretaceous structural elements is given with special emphasis on differences relative to the description by vejbæk (1986). the spatial relationships between these elements (fig. 7) are depicted on the cromer knoll group isochore map (fig. 8). the maps of the early cretaceous geology are included here to emphasise the differences between the structural elements that were active during the late jurassic – earliest cretaceous and those that came into existence during the deposition of the cromer knoll group. the terminology follows that of vejbæk (1986) where reference to other studies is not indicated (table 1). the ål basin (gowers et al. 1993) and the outer rough basin (gowers & sæbøe 1985) are located in the western part of the mapped area and separate the mid north sea high from the inge high and the mads high. these two basins mainly developed during the cretaceous, as pointed out by møller (1986), and are thus only included on the map of early cretaceous structural elements. the ål basin was referred to as the grensen nose by møller (1986) and vejbæk (1986), but this term was originally applied to a structural spur protruding from the mid north sea high (gowers et al. 1993), and is mainly located in norwegian waters. the heno plateau, where lower cretaceous sediments are thin or absent, grades northwards into the nw–se-trending depocentres of the feda graben and the gertrud graben separated by the narrow gert ridge. the heno plateau is bounded to the east by the arne–elin graben. the term lindesnes ridge was applied by gowers & sæbøe (1985) to a late cretaceous – palaeogene inversion structure in the feda graben. the skrubbe fault (gowers & sæbøe 1985) is a complex fault zone separating the ål basin from the feda graben, north of the inge high. the iris basin and gulnare basin (new names) are two separate depocentres that developed in early cretaceous times in the northern parts of the late jurassic tail end graben. salt withdrawal is believed to have played a role in the formation of these depocentres (korstgård et al. 1993); the iris basin is only separated from the gertrud graben by the north arne salt diapir. to the north, the mandal high was an elevated feature and probably experienced erosion during early cretaceous times (see above). the east–west-trending pollerne ridge separates these basins from the roar basin (new name) farther to the south. all three new basin names are taken from local wells. the cromer knoll group is thinly developed in the salt dome province to the south. conclusions the danish central graben is part of the mainly late jurassic complex of grabens which form the north sea central graben. the tail end graben is the dominant late jurassic structural feature in the danish central graben. up to 3600 m of upper jurassic – lowermost cretaceous sediments accumulated in the tail end graben, where deposits in excess of 2000 m thick are encountered for about 90 km along the nw–se-trending coffee soil fault. the rosa basin (new name) is a narrow, north–south-trending basin extending from the south-western part of the tail end graben. the tail end graben ceased to exist as a coherent structural element during the early cretaceous and developed into three separate depocentres: the iris, gulnare and roar basins (new names). the early cretaceous saw a shift from subsidence focused along the coffee soil fault during the late jurassic to a more even distribution of minor basins within the danish central graben. to the west, the ål basin (formerly termed the grensen nose by møller 1986) and the outer rough basin mainly developed during the early cretaceous. the thickness of the cromer knoll group ranges from 700 to 1100 m in the depocentres. 244 acknowledgements the maps are part of the results of an integrated mapping project sponsored by the amoco 3rd round group, comprising amoco denmark, fls-energy, denerco, and dopas. regional mapping of the danish central graben has been undertaken at the survey in the course of the last 20 years. hence many of our former and present colleagues have contributed significantly to the maps presented here: ioannis abatzis, morten sparre andersen, torben bidstrup, carsten langtofte, kurt milthers, carlos møller, jens jørgen møller, kurt damtoft poulsen, erik skovbjerg rasmussen and ole valdemar vejbæk. comments from the referees, g.m. sykes and j.a. korstgård, and the editors significantly improved the manuscript. references andersen, c., olsen, j.c., michelsen, o. & nygaard, e. 1982: structural outline and development. in: michelsen, o. (ed.): geology of the danish central graben. danmarks geologiske undersøgelse serie b 8, 9–26. andsbjerg, j. & dybkjær, k. 2003: sequence stratigraphy of the jurassic of the danish central graben. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 265–300 (this volume). britze, p., japsen, p. & andersen, c. 1995a: geological map of denmark, 1:200 000. the danish central graben. ‘top chalk’ and the post chalk group (two-way traveltime, depth and interval velocity). danmarks geologiske undersøgelse kortserie 47, 7 pp., 3 maps. britze, p., japsen, p. & andersen, c. 1995b: geological map of denmark, 1:200 000. the danish central graben. ‘base chalk’ and the chalk group (two-way traveltime and depth, interval velocity and isochore). danmarks geologiske undersøgelse kortserie 48, 7 pp., 4 maps. britze, p., japsen, p. & andersen, c. 1995c: geological map of denmark, 1:200 000. the danish central graben. ‘base cretaceous’ and the cromer knoll group (two-way traveltime and depth, interval velocity and isochore). danmarks geologiske undersøgelse kortserie 49, 7 pp., 4 maps. britze, p., japsen, p. & andersen, c. 1995d: geological map of denmark, 1:200 000. the danish central graben. ‘base upper jurassic’ and the upper jurassic (two-way traveltime and depth, interval velocity and isochore). danmarks geologiske undersøgelse kortserie 50, 7 pp., 4 maps. damtoft, k., nielsen, l.h., johannessen, p.n., thomsen, e. & andersen, p.r. 1992: hydrocarbon plays of the danish central trough. in: spencer, a.m. (ed.): generation, accumulation and production of europe’s hydrocarbons ii. european association of petroleum geoscientists special publication 2, 35–58. deegan, c.e. & scull, b.j. 1977: a standard lithostratigraphic nomenclature for the central and northern north sea. institute of geological sciences report 77/25, 36 pp. london: her majesty’s stationery office. gowers, m.b. & sæbøe, a. 1985: on the structural evolution of the central trough in the norwegian and danish sectors of the north sea. marine and petroleum geology 2, 298–318. gowers, m.b., holtar, e. & swensson, e. 1993: the structure of the norwegian central trough (central graben area). in: parker, j.r. (ed.): petroleum geology of northwest europe: proceedings of the 4th conference, 1245–1254. london: geological society. ineson, j.r., bojesen-koefoed, j.a., dybkjær, k. & nielsen, l.h. 2003: volgian–ryazanian ‘hot shales’ of the bo member (farsund formation) in the danish central graben, north sea: stratigraphy, facies and geochemistry. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 403–436 (this volume). japsen, p. 1993: influence of lithology and neogene uplift on seismic velocities in denmark: implications for depth conversion of maps. american association of petroleum geologists bulletin 77, 194–211. japsen, p. 1994: retarded compaction due to overpressure deduced from a seismic velocity/depth conversion study in the danish central trough, north sea. marine and petroleum geology 11, 715–733. jensen, t.f., holm, l., frandsen, n. & michelsen, o. 1986: jurassic – lower cretaceous lithostratigraphic nomenclature for the danish central trough. danmarks geologiske undersøgelse serie a 12, 65 pp. johannessen, p.n. 2003: sedimentology and sequence stratigraphy of paralic and shallow marine upper jurassic sandstones in the northern danish central graben. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 367–402 (this volume). johannessen, p.n. & andsbjerg, j. 1993: middle to late jurassic basin evolution and sandstone reservoir distribution in the danish central trough. in: parker, j.r. (ed.): petroleum geology of northwest europe: proceedings of the 4th conference, 271–283. london: geological society. korstgård, j.a., lerche, i., mogensen, t.e. & thomsen, r.o. 1993: salt and fault interactions in the north-eastern danish central graben: observations and inferences. bulletin of the geological society of denmark 40, 197–255. kristensen, l., dons, t., maver, k.g. & schiøler, p. 1995: a multidisciplinary approach to reservoir subdivision of the maastrichtian chalk in the dan field, danish north sea. american association of petroleum geologists bulletin 79, 1650–1660. lieberkind, k., bang, i., mikkelsen, n. & nygaard, e. 1982: late cretaceous and danian limestone. in: michelsen, o. (ed.): geology of the danish central graben. danmarks geologiske undersøgelse serie b 8, 49–62. michelsen, o., nielsen, l.h., johannessen, p.n., andsbjerg, j. & surlyk, f. 2003: jurassic lithostratigraphy and stratigraphic 245 development onshore and offshore denmark. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 147–216 (this volume). møller, j.j. 1986: seismic structural mapping of the middle and upper jurassic in the danish central trough. danmarks geologiske undersøgelse serie a 13, 37 pp. møller, j.j. & rasmussen, e.s. 2003: middle jurassic – early cretaceous rifting of the danish central graben. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 247–264 (this volume). nielsen, l.h. & japsen, p. 1991: deep wells in denmark 1935–1990. lithostratigraphic subdivision. danmarks geologiske undersøgelse serie a 31, 179 pp. rasmussen, e.s. 1995: structural evolution of the gert–mjølner area. marine and petroleum geology 12, 377–385. rasmussen, l.b. 1978: geological aspects of the danish north sea sector. danmarks geologiske undersøgelse iii. række 44, 85 pp. rønnevik, h.c., van den bosch, w. & bandlien, e.h. 1975: a proposed nomenclature for the main structural features in the norwegian north sea. in: finstad, k.g. & selley, r.c. (coordinators): jurassic northern north sea symposium, stavanger, 28–30 september, 1975. norwegian petroleum society (npf) proceedings. jnns/18, 1–16. sundsbø, g.o. & megson, j.b. 1993: structural styles in the danish central graben. in: parker, j.r. (ed.): petroleum geology of northwest europe: proceedings of the 4th conference, 1255– 1267. london: geological society. vejbæk, o.v. 1986: seismic stratigraphy and tectonic evolution of the lower cretaceous of the danish central trough. danmarks geologiske undersøgelse serie a 11, 46 pp. vejbæk, o.v. 1992: geodynamic modelling of the danish central trough. in: larsen, r.m. et al. (eds): structural and tectonic modelling and its application to petroleum geology. norwegian petroleum society (npf) special publication 1, 1–17. ziegler, p.a. 1990: geological atlas of western and central europe, 2nd edition, 239 pp. amsterdam: elsevier for shell internationale petroleum maatschappij. 246 manuscript received 2 february 1996; revision accepted 22 december 1997. e2019430105-01 between 1993 and 2017, denmark was one of the largest oil exporting countries in europe having gained this position from its share in the highly prolific danish central graben. however, outside the central graben few prospects have been adequately mapped, due to a lack of data in these socalled ‘white areas.’ as such, their potential for hydrocarbon accumulation remains uncertain. this paper presents an update of the prospect and play types in this area outside the danish central graben, east of 6°15́ e longitude (fig. 1), based on results from the last 30 years of exploration activities. the paper is part of a resource assessment made by the geological survey of denmark and greenland (geus) to the danish energy agency (schovsbo & jakobsen 2017) and is an update of a former review of the area made in 1987 (thomsen et al. 1987). the succeeding exploration efforts have not changed the overall low expectation for the play types in the area. here, we show that an uncertain resource is associated with both the zechstein carbonate play in the north german basin and the upper triassic – lower jurassic sandstone and lower palaeozoic shale gas plays in northern jylland. however, questions remain as to the source of hydrocarbons in the western offshore area. specifically, we are unable to confirm (or refute) whether these structures are sourced via long-distance migration of hydrocarbons from the danish central graben. exploration history with the 1995 amendments to the danish subsoil act, an open door procedure (odp) was established for the area east of 6°15́ e (hereafter termed the open door area or oda; fig. 1) as a consequence of diminishing interest and lack of competition between oil companies in this part of the danish sub surface. the flexibility of the odp allowed licences to be awarded without holding an actual licensing round and especially sought to motivate small oil companies with innovative ideas to pursue oil and gas exploration. however, the danish government is currently preparing new legislation that will prohibit oil and gas exploration onshore and in inner danish waters, thus terminating new oil and gas exploration activities within a substantial part of the oda. the exploration activities resulting from the odp in 1996 are briefly summarised in table 1, fig. 1 and below. for a full account of the activities prior to 1996, we refer to thomsen et al. (1987) for the period up to 1983, and danish energy agency (1995) for the period up to 1994. following the implementation of the odp in 1995 and until 2017, 27 licences were awarded resulting in the drilling of 5 wells (erik-1, karlebo-1, felsted-1, løve-1 and vendsyssel-1) and the collection of various geochemical and airborne surveys as well as 6700 km 2d and 688 km2 3d seismic surveys. (table 1; figs 1, 2, 3). except for the lower palaeozoic shale gas play only the established plays have been tested and the introduction of the odp has not led to increased drilling activity in eastern denmark. on the contrary, activities began to stagnate in the 1970s when exploration focus shifted abruptly to the danish central graben, following the success of the a-1 well. play types and concepts thomsen et al. (1987) identified the following plays to be relevant in the danish area outside the danish central graben, namely the cambrian sandstone and rotliegend sandstone, zechstein carbonate, bunter sandstone, rhaetian (triassic)–jurassic sandstone, and upper cretaceous – danian chalk plays (fig. 4). in addition to these plays, unconventional shale gas and oil plays have since been identified and are included in this brief review. here, the description is focused on data acquired from hydrocarbon exploration and/or from geothermal activities that have occurred since the thomsen et al. (1987) review. review of hydrocarbon potential in east denmark following 30 years of exploration activities niels h. schovsbo*1 and finn jakobsen1 geus bulletin is free to individuals and institutions in electronic form. the author(s) retain copyright over the article contents. review article | open access geus bulletin vol 43 | e2019430105 | published online: 17 june 2019 https://doi.org/10.34194/geusb-201943-01-05 https://doi.org/10.34194/geusb-201943-01-05 e2019430105-02 cambrian sandstone play hydrocarbon production from cambrian sandstones sourced from lower palaeozoic shales occur in the baltic sea and onshore poland and lithuania. similar sandstones occur throughout denmark with potential source and seal provided by the alum shale formation (fig. 4a). in denmark, expulsion and migration of hydrocarbons occurred during the palaeozoic caledonian orogeny. thermal modelling based on results from the terne-1 well (drilled in 1985) in kattegat and from the vendsyssel-1 well (drilled in 2015) has, however, highlighted the significant risks related to reservoir quality and retention of hydrocarbons for this play in denmark. this implies that it is highly unlikely that the cambrian sandstones hold significant hydrocarbon resources, as previously stated by thomsen et al. (1987). lower palaeozoic shale gas play thermogenic gas trapped in deeply buried palaeozoic shales represents a new play, explored as part of three licenses (4/09, 1/10, and 2/10; table 1) and in similar shales in sweden and poland (schovsbo & nielsen 2017). although studies carried out by the u.s. geological survey (gautier et al. 2014) and by european geological surveys (zijp et al. 2017) indicate a considerable resource potential in denmark, exploration of this play is still limited, and representative well data and production test data for the shales in denmark are lacking. the vendsyssel-1 exploration well revealed gas in the alum shale formation, however, no test production was carried out, so the commercial potential is still unknown. the well was drilled within a so-called ‘sweet spot’, defined as an area with expected highest gas content (schovsbo et al. 2014). results from this well revealed the alum shale to be 40 m thick, compared to 180 m in terne-1. it also contained much less gas than similar prospective parts of poland, which leads to less favourable expectations for future activities, overall. rotliegend sandstone play permian aeolian sandstones form excellent reservoirs in the netherlands, germany and the uk, and similar facies are present in denmark (thomsen et al. 1987). the play occurs along the northern and southern margins of the ringkøbing–fyn high as well as in the horn graben (fig. 1). the rotliegend play has only been tested by a few wells (e.g. felicia-1, drilled in 1987; and borg-1, drilled in 1988). all wells penetrating the rotliegend are characterised by the lack of hydrocarbon shows, indicating that filling of the structures from palaeozoic sources is not likely due to a lack of mature in situ source rocks or the absence of a migration route from other mature source rocks (beha et al. 2008). open door area well well drilled as part of the open door licence open door licences awarded basement high figure 1. schovsbo jakobsen rosa 1/04 2/99 1/10 2/01 3/07 1/03 1/09 1/03 4/99 5/97 1/02 2/10 1/07 1/05 1/14 1/08 4/09 4/99 1/01 2/07 2/09 1/13 1/99 4/972/05 2/97 3/97 1/97 3/99 1/04 2/99 1/10 2/01 3/07 1/03 1/09 1/03 4/99 5/97 1/02 2/10 1/07 1/05 1/14 1/08 4/09 4/99 1/01 2/07 2/09 1/13 1/99 4/972/05 2/97 3/97 1/97 3/99 central graben north german basin a-1 poland terne-1 kattegat baltic sea sweden germany ringkøbing-fyn hign h orn g raben felicia-1 borg-1 løgumkloster-1/2 løve-1 brøns-1 åbenrå-1 jelling-1 varnæs-1 rødby-1/2 søllested-1 norwegian-danish basin fennoscandian boarder zone ringkøbing-fyn hign luna-1 norway erik-1 karlebo-1 felsted-1 vendsyssel-1 danish central graben north german basin a-1 poland terne-1 kattegat baltic sea sweden germany ringkøbing-fyn hign h orn g raben felicia-1 borg-1 løgumkloster-1/2 løve-1 brøns-1 åbenrå-1 jelling-1 varnæs-1 rødby-1/2 søllested-1 norwegian-danish basin fennoscandian boarder zone ringkøbing-fyn hign luna-1 norway erik-1 karlebo-1 felsted-1 vendsyssel-16°15´e 58°n 57°n 56°n 55°n 10°e5°e 15°e10°e5°e fig. 1. open door licences awarded between 1996–2017 and wells mentioned in the text. for details, see table 1. e2019430105-03 the possibility of long-distance migration of hydrocarbons from the danish central graben into structures in the western part of oda was tested in two licences (1/97 and 2/97). this work included geochemical surveys and reprocessing of existing seismic data. however, the results were inconclusive regarding the hypothesis of long-distance hydrocarbon migration. zechstein carbonate play zechstein carbonates sourced by organic-rich intra (z-1 and z-2) carbonate beds and/or the kupferschiefer occur in northern germany and poland and have been extensively explored within the oda. the presence of oil in the løgumkloster wells, as well as hydrocarbon shows in the brøns-1 and åbenrå-1 wells, confirm the presence of a zechstein play along the southern edge of the ringkøbing–fyn high. eight licences focusing on the zechstein carbonates play south (1/01, 1/05, 1/07, 1/08 and 1/14) and north (2/07, 1/09 and 2/09) of the ringkøbing–fyn high have been awarded. the licensing work included collection of seismic data, geochemical analysis of samples, and drilling of the felsted-1 and løve-1 wells. hydrocarbons were found in zechstein carbonate in the felsted-1 well, but the gas contained open door proceduresole consession 1960 1970 1980 1990 2000 2010 2020 ex pl or at io n w el ls 0 5 10 15 20 32 wells drilled 1935–1960 1 2 3 4 5 6 7 8 1 fe lst ed -1 er ik -1 ka rle bo -1 ve nd sy ss el -1 lø ve -1 primary target for wells east of 6o15´e rhaetian–jurassic bunter licencing round number zechstein pre zechstein exploration wells outside danish central graben exploration wells total figure 2. schovsbo jakobsen, rosa fig. 2. drilling activities and concession rounds in denmark 1960–2017. based on thomsen et al. (1987) for the period up to 1983. after the open door procedure (odp) was established the licencing rounds included only the area west of 6°15́ e. licences based on the eighth licence round have not yet been awarded (application deadline 1 february 2019). !( !(!( !( !(!( !(!( kvols-2,3 margretheholmen-1,2 stenlille-19,20 sønderborg-1,2 karlebo-1 vendsyssel-1 felsted-1 løve-1 !( figure 3. schovsbo jakobsen rosa erik-1erik-1 3d seimic after 1996 basement high open door area well geothermal/storage drilled after 1996 well drilled as part of the open door licence oil and gas related 2d seimic after 1995: aeromagnetic/gravimetry geothermal/storage related 58°n 57°n 56°n 55°n 10°e 15°e10°e fig. 3. seismic data, aeromagnetic/gravity survey and wells drilled after 1995. e2019430105-04 more than 90% n2. the results of the exploration work confirmed the presence of hydrocarbons in the zechstein carbonates but have not proven the presence of prospects of an economical size. new information on the presence of source rocks has not been established. on the northern flank of the ringkøbing–fyn high the jelling-1 well (drilled in 1992) penetrated only 3 m into zechstein carbonates and no source rocks were identified. also, the løve-1 well (drilled in 2011) was dry with no indication of source rocks. the exploration in the northern part of the ringkøbing–fyn high has thus not provided any new or positive indications of good reservoir intervals nor the presence of source rocks. bunter sandstone play good to excellent quality lower triassic reservoirs are widespread in denmark (mathiesen et al. 2010; kristensen et al. 2016). the presence of n2 gas in bunter in the tønder structure confirmed the possibility of a triassic play in the area, but it is dependent on the presence of structural closures and migration from mature source rocks. drilling on salt structures (varnæs-1, rødby-1, -2, and søllested-1, drilled 1952– 1982) found no traces of hydrocarbons. therefore, the triassic play in the north german basin is uncertain with the lack of source rocks being the biggest risk factor. in the horn graben, two licences (4/97 and 1/99) with a bunter sandstone play as the target have been awarded. one dry well was drilled (erik-1, 2001), which confirmed the results from previous wells in the area and verified the good reservoir properties of this interval. a basin modelling and structural development study of the area (beha et al. 2008) revealed the lack of mature source rocks as the biggest determining factor for the presence of hydrocarbons at this location. rhaetian–jurassic sandstone play until 1987, the bulk of the drilling activity in the oda was directed towards the rhaetian–jurassic sandstone reservoirs in the norwegian–danish basin (fig. 4b). based on the available source rock analysis and maturity modelling it appears that only minor amounts of hydrocarbons may have been generated and expelled from lower jurassic shales in deep rim synclines around the salt diapirs (petersen et al. 2008). the rhaetian–jurassic sandstone play was investigated in licences 2/99 and 1/04 (fig. 1). as part of the work programme for license 1/04, a comprehensive 2d seismic survey was carried out along with source rock screening of existing wells. the evaluation report for these wells stated that there are no mature source rocks in the area. mature lower jurassic source rocks (with respect to oil generation) were, however, present in the farsund basin located in the norwegian sector (fig. 4b). but none of the investigations associated with the licences suggested mature hydrocarbons in rim palaeogene periodma a era cretaceous jurassic triassic permian carboniferous devonian silurian ordovician cambrian precambrian pa lae oz oi c m es oz oi c c en . 66 145 201 252 299 359 419 444 485 541 lithostratigraphic units chalk gp zechstein gp haldager fm fjerritslev fm gassum fm bunter sst lower carboniferous ? westphalien ? rotliegend gp source rock reservoir lower cambrian sst. alum shale fm figure 4a. schovsbo, jakobsen rosa fars und basin figure 4b schovsbo jakobsen, rosa rhaetian-jurassic play salt diapirs/pillows potential jurassic kitchens zechstein carbonate play 58°n 57°n 56°n 55°n 10°e 10°e b fig. 4. a: stratigraphical column with petroleum system elements for the norwegian-danish basin. b: zechstein and rhaetian–jurassic play map. fm: formation, gp: group, sst: sandstone. for map legend, see fig. 1. based on thomsen et al. (1987) and petersen et al. (2008). e2019430105-05 synclines around salt structures, nor did they verify filling of structures by migrating hydrocarbons from mature source rocks in the norwegian sector. rhaetian–jurassic sandstones charged from local lower jurassic kitchens – areas where source rocks have reached suitable temperature and pressure to generate hydrocarbons – located in deep rim synclines, were targets in licence 3/07 in western jylland (fig. 1). the work programme did not result in any drilling activities, since only a limited hydrocarbon generative potential – in terms of maturity and quality – of the source rock is expected to be present in the area. upper cretaceous – danian chalk play the chalk play is highly prolific in the danish central graben where it is the single most important reservoir. in this play the maastrichtian–danian chalks are sourced by upper jurassic to lowermost cretaceous marine shales and sealed by paleogene shales or intra-chalk tight zones. however, outside the danish central graben area, maturity modelling indicates that these jurassic shales may not have reached oil-maturity. therefore, the play in the oda has to rely on long-distance migration from the danish central graben as tested by the two licences (1/97 and 2/97). however, the investigations did not yield a final conclusion on the existence of long-distance migration from the danish central graben into the oda. conclusions the danish area east of 6°15´e contains several plays including the lower palaeozoic unconventional gas, zechstein carbonate in the north german basin and the rhaetian (triassic) – lower jurassic sandstone in northern jylland. however, exploration activities have not been able to demonstrate the existence of commercial quantities of hydrocarbons. for the zechstein carbonate play, all identified structures appear to be of limited size, the reservoir quality is generally poor, and the presence of source rocks is uncertain. the rhaetian–jurassic sandstone plays depend on local kitchens developed around salt diapirs. current data suggest that kitchen areas are table 1. summary of hydrocarbon exploration activities in the open door area (oda) after the establishment of the open door procedure (odp) in 1996. license 1/97 2/97 3/97 4/97 5/97 1/99 2/99 3/99 4/99 1/01 2/01 1/02 1/03 1/04 1/05 2/05 1/07 2/07 3/07 1/08 1/09 2/09 4/09 1/10 2/10 1/13 1/14 operator agip amerada amerada mærsk odin agip gustavson anschutz amerada minijos nafta sterling tethys tethys dong wexco elko geo-center-nord jordan dong danica danica danica schuepbach total total nikoil jutland petroleum new data acquired seismic: 81 km 2d, 347 km² 3d seismic: 4014 km 2d; well: erik-1 geochemical survey seismic: 213 km 2d; geochemistry and aeromagnetic survey seismic: 165 km 2d, 225 km² 3d geochemical survey geochemical survey well: karlebo-1 geochemical survey seismic: 1661 km 2d seismic: 37 km 2d, 105 km² 3d; well: felsted-1; geochemical survey well: luna-1 with licence 1/11 seismic: 20 km 2d well: løve-1 seismic: 50 km 2d; seabed cores; geochemical survey seismic: 38.5 km 2d; geochemical survey seismic: 2d, 3d; geochemical survey seismic: 2d, 3d; geochemical survey well: vendsyssel-1 aeromagnetic and gravimetric survey primary target chalk sourced danish central graben palaeocene sandstones sourced from danish central graben palaeocene sandstones sourced from danish central graben bunter sandsone sourced from carboniferous sources haldager/gassum sandstone sourced from fjerritslev fm bunter sandsone sourced from carboniferous sources haldager/gassum sandstone sourced from fjerritslev fm cretaceous sandstone sourced from jurassic shales zechstein carbonates sourced from z1/z2 basin facies or kupferschiefer mesozoic sandstones sourced from zechstein or lower palaeozoic shales. gassum/haldager sandstone sourced from fjerritslev fm or intra-gassum zechstein carbonates sourced from z1/z2 basin facies or kupferschiefer long-distance migration from danish central graben zechstein carbonates sourced from z1/z2 basin facies or kupferschiefer zechstein carbonates haldager sandstone sourced from fjerritslev fm zechstein carbonates zechstein carbonates lower palaeozoic shale gas lower palaeozoic shale gas lower palaeozoic shale gas zechstein carbonates zechstein carbonates turbidites sourced from stinkkalk/shale a complete record of data collected as part of each license is provided in schovsbo & jakobsen (2017). e2019430105-06 *corresponding author: niels h. schovsbo | e-mail: nsc@geus.dk 1 geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350, copenhagen k, denmark. limited in extent. the palaeozoic shales appear to be thinner than expected and to include less gas than similar prospective shales in poland. exploration activities during the past 30 years (including the odp) confirm the presence of the previously identified uncertain plays in the oda. the odp has not identified new conventional plays and has not proved or disproved the major unknowns in the area, namely the presence of mature source rock. instead, data gathered for these activities have highlighted the excellent reservoir properties of rhaetic–jurassic and bunter sandstones as particularly appropriate geothermal reservoirs. acknowledgments this is a part of an ongoing eu horizon 2020 geoera project (the garah, h2020 grant #731166 lead by geus). references beha, a., thomsen, r.o. & littke, r. 2008: thermal history, hydrocarbon generation and migration in the horn graben in the danish north sea: a 2d basin modelling study. international journal of earth sciences 97, 1087–1100. https://doi.org/10.1007/s00531-007-0247-2 danish energy agency 1995: rapport om efterforskningspotentialet i øst-danmark (including english summaries). arbejdsgruppen vedr. vurdering af det østdanske efterforskningspotentiale, 50 p. unpublished report, danish energy agency, geological survey of denmark and greenland & dopas (in archives of geological survey of denmark and greenland, geus report file no. 13086). gautier, d.l., schovsbo, n.h. & nielsen, a.t. 2014: resource potential of the alum shale in denmark. proceedings of the 2nd unconventional resources technology conference. denver, colorado, usa, 25–27 august, 2014. https://dx.doi.org/10.15530/urtec-2014-1931754 kristensen, l., hjuler, m.l., frykman, p., olivarius, m., weibel, r., nielsen, l.h. & mathiesen, a. 2016: pre-drilling assessments of average porosity and permeability in the geothermal reservoirs of the danish area. geothermal energy 4, 27 pp. https://doi.org/10.1186/s40517016-0048-6 mathiesen, a., nielsen, l.h. & bidstrup, t. 2010: identifying potential geothermal reservoirs in denmark. geological survey of denmark and greenland bulletin 20, 19–22. petersen, h.i., nielsen, l.h., bojesen-koefoed, a., mathiesen, a., kristensen, l. & dalhoff, f. 2008: evaluation of the quality, thermal maturity and distribution of potential source rocks in the danish part of the norwegian–danish basin. geological survey of denmark and greenland bulletin 16, 66 pp. schovsbo, n.h. & jakobsen, f. 2017: kulbrinteefterforskningspotentialet i åbendørområdet samt betydningen af hydraulisk frakturering for udnyttelsen af kulbrinteressourcer og af geotermi-ressourcer, 18 pp. unpublished report (in archives of geological survey of denmark and greenland, geus report file no.34280). schovsbo, n.h. & nielsen, a.t. 2017: generation and origin of natural gas in lower palaeozoic shales from southern sweden. geological survey of denmark and greenland bulletin 38, 37–40. schovsbo, n.h., nielsen, a.t. & gautier, d.l. 2014: the lower palaeozoic shale gas play in denmark. geological survey of denmark and greenland bulletin 31, 19–22. thomsen, e., damtoft, k. & andersen, c. 1987: hydrocarbon plays in denmark outside the central trough. in: brooks, j. & glennie, k.w. (eds): petroleum geology of north west europe: proceedings of the 3rd conference on petroleum geology of north west europe, barbican centre, london, 26–29 october, 1986. geological society of london, uk. zijp, m., nelskamp, s., schovsbo, n.h., tougaard, l. & bocin-dumitriu, a. 2017: resource estimation of eighty-two european shale formations. proceedings of the 5th unconventional resources technology conference, austin, texas, usa, 24–26 july, 2017. https://doi.org/10.15530/ urtec-2017-2686270 how to cite schovsbo, n.h., jakobsen, f. 2019: review of hydrocarbon potential in east denmark following 30 years of exploration. geological survey of denmark and greenland bulletin 43, e2019430105. https://doi.org/10.34194/geusb-201943-01-05 mailto:nsc@geus.dk https://doi.org/10.1007/s00531-007-0247-2 https://dx.doi.org/10.15530/urtec-2014-1931754 https://doi.org/10.1186/s40517-016-0048-6 https://doi.org/10.1186/s40517-016-0048-6 https://doi.org/10.15530/urtec-2017-2686270 https://doi.org/10.15530/urtec-2017-2686270 https://doi.org/10.34194/geusb-201943-01-05 geological survey of denmark and greenland bulletin 31, 2014, 79-82 79 outlet glacier dynamics and bathymetry at upernavik isstrøm and upernavik isfjord, north-west greenland camilla s. andresen, kristian k. kjeldsen, benjamin harden, niels nørgaard-pedersen and kurt h. kjær during the past decades, the greenland ice sheet has experienced a marked increase in mass loss resulting in an increased contribution to global sea-level rise. the three largest outlet glaciers in greenland have increased their discharge, accelerated, thinned and retreated between 1996 and 2005. after 2005 most of them have slowed down again although not to previous levels. geodetic observations suggest that rapid increase in mass loss from the north-western part of the ice sheet occurred during 2005–2010 (kjeldsen et al. 2013). warming of the subsurface water masses off greenland may have triggered the acceleration of outlet glaciers from the ice sheet (straneo & heimbach 2013). the north atlantic subpolar gyre, which transports water to south-east and west greenland via the warm irminger current, warmed in the mid-1990s. increased inflow of warm subpolar waters likely led to increased submarine melting of tidewater glaciers. climate, glacier configuration and fjord bathymetry play fundamental roles for outlet glacier dynamics and thus knowledge of these parameters is warranted. in particular, the bathymetry of a fjord gives important information about the exchange between fjord waters close to marine-terminating glaciers and the shelf and ocean. however, only sparse bathymetric data are available for the majority of fjords in greenland. the international bathymetry chart for the arc© 2014 geus. geological survey of denmark and greenland bulletin 31, 79–82. open access: www.geus.dk/publications/bull a 73°n glacier 1 glacier 2 glacier 3 glacier 4 55°w 5 km potential temperature (°c) salinity (psu) 1 2 3 32 33 34 35 100 200 300 400 500 600 700 d ep th (m ) 73°n 56°w upernavik isstrøm b upernavik isfjord 10–100 101–200 201–300 301–400 401–500 501–600 601–700 701–800 801–900 901–1000 1001–1100 water depth depth (m) greenland hg mb fig. 1. a: landsat 8 satellite image from august 2013 of the upernavik isfjord region with recorded water depths in the fjord. b: temperature and salinity profiles at a mid-fjord site (white diamond). data acquired in 2013 by oceanographers from woods hole oceanographic institution. mb: melville bugt. hg: helheimgletscher. 8080 tic ocean (ibcao) does not provide adequate data for the fjords and gives the impression that water depths in fjords are typically <200 m. here we present the first detailed bathymetric data from upernavik isfjord in north-west greenland, which were obtained during a cruise led by the geological survey of denmark and greenland in august 2013. the purpose of the cruise was to retrieve sediment cores, collect hydrographic data and map the bathymetry of the fjord. in this paper, we also estimate retreat rates of the upernavik isstrøm since 1849 and evaluate them in the context of climate variability, glacier setting and fjord bathymetry. upernavik isstrøm and upernavik isfjord upernavik isstrøm consists of four main ice streams (glaciers 1–4, fig. 1) that had a total mass loss of 53.5 ± 12.8 gt during the period 2005–2010 (khan et al. 2013). the four glaciers terminate in the c. 80 km long upernavik isfjord. glaciers 1 and 2 are the most productive and the fjord in front of these glaciers is packed with icebergs throughout the year. the bathymetric data (fig. 1) show that most of the fjord is over 900 m deep, but water depths of 600–800 m are found near its head. due to ice conditions, water depths could only be measured near glacier 4 where there is an area with water depths around 200 m. local fishermen report water depths of 600–700 m at a distance of c. 5–10 km from the fronts of glaciers 1–3. the survey ended slightly west of the fjord mouth and there was no indication of a sill. the hydrographic measurements show a c. 2°c warm, lowsalinity, 50 m thick surface layer. from 50 to 150 m cold polar water with a temperature of 0.5–1.5°c is found, and below this the water gradually warms from 1 to 3°c and becomes more saline, which shows that atlantic water penetrates into the fjord (fig. 1). this also indicates that there is no shallow sill at the entrance to the fjord. radar-based surveys suggest grounding line depths of 400–700 m for glaciers 1–3 and 100 m for glacier 4 (morlighem et al. 2014). this suggests that atlantic water comes into contact with the fronts of glaciers 1–3, whereas the front of glacier 4 is in contact with polar water. this finding has implications for understanding the history of glacier retreat. glacier retreat and climate change frontal positions of upernavik isstrøm were compiled for the period from 1849 to 1953 by weidick (1958; fig. 2). along with satellite images from 1966 and 1976 (this study) and more continuously since 1985 (khan et al. 2013) these compilations provide a 150 years long record of glacier retreat since the little ice age maximum position. using the centre flow-line we estimate average annual retreat rates between glacier margin positions (fig. 3c). from the little ice age until c. 1931, the four glaciers were merged into one and retreated relatively slowly. the frontal retreat rate accelerated around 1931 and glaciers 1 and 2 and glaciers 3 and 4 started to split into two separate arms, and after 1946 glaciers 3 and 4 were decoupled from each other. the relatively high retreat rates lasted until the mid-1940s and were followed by lower retreat rates. glaciers 1 and 2 decoupled from each other after 1966. three subsequent episodes of increased retreat rates are seen: (1) between 1966 and 1985 (glaciers 1, 2 and 4), (2) late 1990s (glaciers 1, 2 and 4) and (3) 2005–2009 (all glaciers). during the latter period, the retreat rate of glacier 1 was exceptionally high. the temporal resolution of the data does not allow us to detect earlier similar rapid retreat events. in a recent study, data on frontal changes, thinning, and glacier velocity since 1985 were obtained from aerial photos and sat73°n 54°w 54°w55°w 55°w 5 km upernavik isfjord 1849 1886 1931 1937 1942 1946 1949 1953 1966 1976 1985 1991 1996 2000 2005 2010 fig. 2. glacier frontal positions based on maps from historical expeditions, aerial photographs and satellite images (weidick 1958; khan et al. 2013), supplemented with a corona satellite image from 1966 and a landsat mss image from 1976. the black and white lines show the tracks used for calculating single-point distances from the glacier margin. 81 ellite data and used to estimate dynamic mass loss (khan et al. 2013). glacier 4 experienced a marked mass loss episode, including >100 m thinning, prior to 1991. at the same time no marked changes were recorded for glaciers 1–3. in contrast, between 2005 and 2009, glacier 1 sped up, retreated and thinned markedly, whereas glaciers 2–4 were relatively stable. kjær et al. (2012) documented two events of dynamic mass loss along the melville bugt coast but also noted a spatially variable pattern in the magnitude of these events. the extended retreat data presented here may indicate that the marked dynamic mass loss prior to 1991 of glacier 4 was preceded by an even more marked event between 1966 and 1985 affecting not only glacier 4 but also glaciers 1 and 2. care should be taken in comparing magnitude and timing of retreat rates from merged and decoupled glaciers due to their different and temporally variable tributary sizes, glacier tongue widths and flow rates. moreover, the estimated frontal positions are based on single points and may be subject to seasonal fluctuations and are thus only approximate. however, the timing of accelerated retreat rates in the different glaciers is fairly synchronous over inter-annual time scales. the onset of increased retreat rates in the 1930s of glaciers 1 + 2 and 3 + 4 as well as the increased retreat rates of glaciers 1 + 2 around 1966 may have been initiated by their decoupling from the merged glacier and glaciers 1 + 2 at this time. this would suggest that topographical constraints are important for retreat rates. alternatively, a common climate change may have forced accelerated retreat and in this way caused the decoupling of glaciers. the marked retreat episodes in 1931–1946 (of the merged glacier), in the late 1990s (glaciers 1, 2, 4) and in 2005–2009 (all glaciers, but most markedly glacier 1) occurred at times of marked warming near upernavik and entire greenland (chylek et al. 2006). marked retreat from 1930 to the 1940s has also been documented elsewhere in greenland, specifically for helheimgletscher in south-east greenland (andresen et al. 2012; figs 1, 3d) and has been ascribed to variations in the atlantic multi-decadal oscillation (amo). the amo is a mode of variability with its main expression in sea-surface temperatures in the north atlantic ocean and influencing circum-atlantic climate, including coastal greenland. the amo has a periodicity of c. 60 years, and a positive amo fig. 3. a: atlantic multi-decadal oscillation (amo) index (schlesinger & ramankutty 1994). b: annual average air temperature for upernavik (data from the danish meteorological institute). c: calculated changes in rate (m yr–1). negative values and colour-filled boxes: glacier retreat. positive values and white boxes: glacier advance. episodes characterised by increased retreat rates by upernavik are highlighted with yellow boxes. d: marine sediment-based proxy data from south-east greenland. relative variability in calving from helheimgletscher based on sand fluxes (andresen et al. 2012) and shelf sea-surface temperatures based on analyses of core er07 from sermilik fjord (andresen et al. 2013). yellow boxes highlight warm episodes with increased dynamic mass loss in south-east greenland. 1860 1880 1900 1920 1940 1960 1980 2000 year amo+ amo– glaciers 1–4 merged glacier 1 glacier 2 glacier 4 glacier 3 1+2 merged se greenland 1+2 merged 3+4 merged 3+4 merged glaciers 1–4 merged glaciers 1–4 merged glaciers 1–4 merged upernavik isstrøm –2293 g lac ie r r et re at ra te (m /y ea r) 12 calving helheimgletscher te m p. (° c ) a b c d –10 –6 –2 te m pe ra tu re (° c ) 0 –400 –800 400 0 –400 –800 400 0 –400 –800 0 –500 8 10 8282 index is linked with higher sea-surface temperatures. the accelerated retreat episodes of upernavik isstrøm from 1930 to the 1940s, late 1990s and 2005–2009 could be associated with warming of subsurface waters during periods with positive amo indexes (fig. 3). however, since air temperatures co-vary with the amo index on multi-decadal timescales (figs 3a, b) it is not possible to differentiate directly between influence from increased air versus water temperatures on mechanisms that could lead to retreat of the glacier margin. as noted the increased retreat rates of glaciers 1 and 2 between 1966 and 1985 may have been triggered by changed topographical constraints such as loss of pinning points as the combined glacier front widened considerably (figs 2, 3c). however, due to the synchronous, high retreat rates of glacier 4 within its own tributary, as well as continued high retreat rates long after decoupling of glaciers 1 and 2, we speculate that climate forcing is also involved. the increased retreat differs from the other retreat episodes because it occurred during a negative amo index. however, even though average sea-surface temperatures in the source region of atlantic waters were generally low between the early 1960s and the mid-1990s and annual air temperatures in upernavik decreased slightly, air temperatures increased between the early 1970s and 1980. interestingly, sediment-based proxy glacier and ocean data document a marked warming of atlantic waters in south-east greenland at this time along with a marked increase in calving from helheimgletscher (fig. 3d). the concurrency between this climate warming and glacier instability in south-east greenland supports the theory that the increased retreat rates of glaciers 1, 2 and 4 between 1976 and 1985 were forced by climate warming. the front of glacier 4 has been located in a 200 m shallow area since the 1950s (fig. 1) and the front of this glacier is not in direct contact with the deeper warm subsurface layer in the fjord. although warming or increased thickness of the atlantic water layer may increase the temperature of the polar water layer, we suggest that the accelerated retreat rates of glacier 4, including the 1985–1991 episode of marked thinning (khan et al. 2013), may be linked with increased air temperatures. it has been suggested that meltwater percolating down the glacier being released as subglacial discharge may have a considerable influence on the submarine melt rates and thus glacier stabilisation. we emphasise that the retreat record is not fully representative of glacier changes. not only are the determined frontal positions rather sporadic and do not represent a continuous record of change, but in addition we need to assess thinning and flow-rate changes to obtain a more comprehensive picture of glacier changes since the little ice age. one way to overcome this will be to analyse sediment cores to obtain a continuous proxy record of calving variability. this will add information on dynamic changes and can subsequently be linked to digital elevation models and mass-balance modelling to estimate mass loss. our results also show that bathymetric conditions may partly explain asynchronous glacier responses to climatic warming. it is therefore important to incorporate bathymetric data when trying to understand and predict outlet-glacier behaviour. hopefully future campaigns to collect and pool bathymetric data will provide improved bathymetric maps of the greenland fjords. acknowledgements the ‘upernavik glacier project’ is funded by geocenter danmark. we thank arctic station on disko for the use of the vessel porsild and hans karl petersen from upernavik who was pilot during the survey of glacier 4 and provided water-depth information by glaciers 1–3. references andresen, c.s. et al. 2012: rapid response of helheim glacier in greenland to climate variability over the past century. nature geoscience 5, 37–41. andresen, c.s., sicre, m.-a., straneo, f., sutherland, d.a., schmith, t., ribergaard, m.h., kuijpers, a. & lloyd, j.m. 2013: a 100-year record of alkenone-derived sst changes by southeast greenland. continental shelf research 71, 45–51. chylek, p., dubey, m.k. & lesins, g. 2006: greenland warming of 1920– 1930 and 1995–2005. geophysical research letters 33, l11707. khan, s.a., et al. 2013: recurring dynamically induced thinning during 1985 to 2010 on upernavik isstrøm, west greenland. journal of geophysical research: earth surface 118, 111–121. kjær, k.h. et al. 2012: aerial photographs reveal late-20th-century dynamic ice loss in northwestern greenland. science 337, 569–573. kjeldsen, k.k., et al. 2013: improved ice loss estimate of the northwestern greenland ice sheet. journal of geophysical research: solid earth 118, 698–708. morlighem, m., rignot, e., mouginot, j., seroussi, h.& larour, e. 2014: deeply incised submarine glacial valleys  beneath the greenland ice sheet. nature geoscience 7, 418–422. schlesinger, m.e. & ramankutty, n. 1994: an oscillation in the global climate system of period 65–70 years. nature 367, 723–726. straneo, f. & heimbach, p. 2013: north atlantic warming and the retreat of greenland’s outlet glaciers. nature 504, 36–43. weidick, a. 1958: frontal variations of upernaviks isstrøm in the last 100 years. meddelelser fra dansk geologisk forening 14, 52–60. authors’ addresses c.s.a. & n.n.p., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: csa@geus.dk k.k.k. & k.h.k., natural history museum of denmark, university of copenhagen, øster voldgade 5-7, dk-1350 copenhagen k, denmark. b.h., woods hole oceanographic institution, woods hole, ma 02543, usa. geological survey of denmark and greenland bulletin 1, 9-20 9 the jurassic of denmark is mainly known from subsurface data including numerous boreholes and a dense net of seismic lines in the offshore areas. exceptions to this include the island of bornholm in the baltic sea and adjacent areas of skåne, southern sweden, where the jurassic is exposed in many small outcrops. the jurassic of greenland, in contrast, can be studied in extensive outcrops and the succession forms the walls and tops of mountains and plateaus over wide areas. the two regions were once part of the same large-scale system of extensional basins in the northwest european – north atlantic region (fig. 1), but are today located on two different plates separated by a thousand kilometres of ocean. the aim of this introductory paper is to compare and contrast the stratigraphic evolution of the two regions based primarily on the detailed studies included in this book; the focus is on the timing and nature of tectonic events and on the overall stratigraphic trends. the goal is to provide a broad evolutionary framework for the jurassic of denmark and greenland to set the scene for the succeeding papers. for comprehensive reviews of the north atlantic mesozoic rift system, the the jurassic of denmark and greenland: key elements in the reconstruction of the north atlantic jurassic rift system finn surlyk and jon r. ineson the jurassic succession of denmark is largely confined to the subsurface with the exception of exposures on the island of bornholm in the baltic sea. in east greenland, in contrast, the jurassic is extensively exposed. comparison of basin evolution in the two regions, which now occur on two separate plates, thus relies on highly different datasets. it is possible nevertheless to construct an integrated picture allowing testing of hypotheses concerning basin evolution, regional uplift, onset and climax of rifting, relative versus eustatic sea-level changes and sequence stratigraphic subdivision and correlation. on a smaller scale, it is possible to compare the signatures of sequence stratigraphic surfaces as seen on well logs, in cores and at outcrop and of sequences recognised and defined on the basis of very different data types. breakdown of the successions into tectonostratigraphic megasequences highlights the high degree of similarity in overall basin evolution and tectonic style. an important difference, however, lies in the timing. major events such as late early – middle jurassic uplift, followed by onset of rifting, basin reorganisation and rift climax were delayed in east greenland relative to the danish region. this has important implications both for regional reconstructions of the rift system and for the understanding and testing of classical sequence stratigraphic concepts involving eustatic versus tectonic controls of basin evolution and stratigraphy. keywords: denmark, greenland, jurassic, correlation, parallel evolution f.s., geological institute, university of copenhagen, geocenter copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: finns@geo.geol.ku.dk j.r.i., geological survey of denmark and greenland, geocenter copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ji@geus.dk geological survey of denmark and greenland bulletin 1, 9–20 (2003) © geus, 2003 reader is referred to ziegler (1988, 1990), doré (1992), doré et al. (1999), skogseid et al. (2000) and brekke et al. (2001). sequence stratigraphy – a conceptual note much of the research presented in this book is based on sequence stratigraphic analysis and it is pertinent in this introduction to comment briefly on the conceptual basis for these studies. as stressed by many workers (e.g. carter et al. 1991; posamentier & james 1993; miall 1997), sequence stratigraphy may be applied in two fundamentally different ways, either involving construction of age models based on correlation with the so-called global cycle chart of haq et al. (1987) or lithology prediction based on the interpretation of cyclicity in the rock record (posamentier & james 1993). these two distinct paradigms were termed the ‘global– eustasy paradigm’ and the ‘complexity paradigm’ by miall & miall (2001). it is significant that few of the authors of the individual studies presented here use the ‘global–eustasy paradigm’ but prefer the ‘complexity’ model which deals with the stratigraphic architecture and predictability of sequences and their components. emphasis is on the recognition, interpretation and dating of key surfaces and on the geometry and environmental nature of successive systems tracts. there is, in contrast, little reference to ‘global cycle charts’ and to the potential use of sea-level curves as dating tools. rather, the ages and significance of key surfaces and derived sea-level curves are used to construct robust genetic stratigraphies, to chart basin evolution and to highlight the importance of timing of tectonic events and pulses of sediment input. this approach is in marked contrast to that adopted by most authors in the compilation by de graciansky et al. (1998) in which the ‘global–eustasy paradigm’ is prevalent. tectonostratigraphy subdivision of the jurassic successions of both regions into tectonostratigraphic packages (sensu surlyk 1991) shows that the tectonic evolution and corresponding stratigraphic signals are broadly similar, although the timing of the transition from pre-rift uplift to onset of rifting and of rift culmination appears to be delayed in east greenland compared to the north sea region. the sedimentary environments and facies are also rather similar for the successive tectonostratigraphic units. the early pre-rift succession (rhaetian–sinemurian) of east greenland is wholly non-marine, however, contrasting with the marine late triassic – early jurassic record of much of the danish region. conversely, the aalenian– callovian early syn-rift sediments of the north sea and the danish basin are more proximal and terrestriallydominated than the correlatives in east greenland which are almost exclusively marine. the investigated time interval includes the uppermost triassic and the lowermost cretaceous in order to cover a complete set of genetically related units. this stratigraphic interval in denmark and greenland can be broadly compared in terms of two megasequences, an upper triassic – middle jurassic pre-rift megasequence and a middle jurassic – lowermost cretaceous syn-rift megasequence, separated by a regional uplift event (fig. 2). detailed correlation between the two regions allowing comparison of short-term sea-level cycles is not yet possible. the pre-rift biostratigraphy is based to a large extent on dinocysts (poulsen & riding 2003, this volume), which are stratigraphically rather long ranging and commonly show different ranges in the two regions. correlation by ammonites can only be done at a few levels, notably in the pliensbachian. dating of the syn-rift succession is based mainly on dinocysts in the north sea – danish basin region but on ammonites supported by dinocysts in east greenland (callomon 2003, this volume; surlyk 2003, this volume). faunal provincialism was strongly developed for much of the middle and late jurassic. pre-rift megasequence in east greenland, the rhaetian – early bajocian time interval was characterised by regional subsidence following rift events in the late permian and early triassic (surlyk 1990, 2003, this volume). the depositional basin was centred over jameson land and stratigraphic units have a more or less basinwide extent and sheet-like geometry, reflecting the relatively uniform subsidence and the absence of major faulting (dam & surlyk 1995, 1998). in general, however, individual units are thickest in the basin centre. a similar development is seen in denmark, the upper triassic – aalenian succession recording a phase of relatively uniform regional subsidence following rift events in the late carboniferous – early permian and the early–middle triassic. the sedimentary record is fragmentary, however, as lower jurassic rocks are missing 10 11 60˚n 55˚n 50˚n 45˚n pa lae ol at itu de 500 km wollaston forland kuhn ø andøy jameson land milne land mid north sea high cg united kingdom denmark db skåne bornholm in tr arif t h igh middle jurassic land normal fault igneous activity inferred structural high deltaic/shallow marine sandstone offshore marine mudstone marine carbonate ? greenland norway sweden baltic shield laurentian shield ringkøbing–fyn high fig. 1. schematic middle jurassic reconstruction showing the regional tectonic elements and jurassic seaways in the north sea region and between greenland and norway. for location of intrabasinal structural elements (danish basin, danish central graben) named in the text, the reader is referred to michelsen et al. (2003, this volume). map based on ziegler (1988, 1990), doré (1992) and surlyk (2003, this volume). cg, central graben; db, danish basin. from large parts of the area, particularly the central graben, due to erosion following early middle jurassic uplift (e.g. ziegler 1990; underhill & partington 1994; andsbjerg et al. 2001; andsbjerg & dybkjær 2003, this volume; nielsen 2003, this volume). the erosional remnants, which are located marginal to and outside the main uplifted areas, are indicative of laterally extensive and sheet-like sedimentary packages, similar to those described from east greenland. this architectural style is recorded, in particular, by the marine lower jurassic fjerritslev formation, which is recognised both in the danish sector of the central graben and in the danish basin (michelsen et al. 2003, this volume). over much of the danish area, the depositional environments and facies are more offshore marine and finer-grained than in the land-locked jameson land basin of east greenland although the paralic successions on bornholm and in skåne display alternating lacustrine, estuarine and shoreface deposits that closely resemble the jameson land succession and show a comparable overall transgressive trend (ahlberg et al. 2003, this volume; frandsen & surlyk 2003, this volume; michelsen et al. 2003, this volume). the pre-rift succession in east greenland can be subdivided into a rhaetian–sinemurian fluvial–lacustrine part and a pliensbachian – early bajocian estuarine – offshore marine part (dam & surlyk 1995, 1998; surlyk 2003, this volume). although a similar gross subdivision, reflecting an overall transgressive trend, can be demonstrated in the lower jurassic of denmark, the timing of the marine inundation of non-marine/paralic settings can only be compared in detail with that of east greenland in the most proximal areas (skagerrak–kattegat platform, skåne, bornholm). here, the rhaetian–hettangian succession was deposited in terrestrial environments succeeded by paralic sinemurian and offshore marine pliensbachian conditions. in the danish central graben and the axial parts of the danish basin, marine conditions were already established in late triassic – earliest jurassic times (fig. 2). a composite regressive event took place in the danish basin in the rhaetian, corresponding broadly to a hiatus in the danish central graben (fig. 2); the subsequent hettangian–sinemurian period was a time of stepwise deepening and expansion of the open marine environment, reaching bornholm in the latest sinemurian (surlyk et al. 1995; nielsen 2003, this volume, fig. 31). hence, inundation of the nonmarine jameson land basin at the sinemurian–pliensbachian boundary was broadly coeval with a long-term maximum transgression in the danish area (fig. 2). in the danish basin, the toarcian records the onset of regression and progressive basin restriction, heralding regional uplift and erosion and the reversion to terrestrial conditions in the middle jurassic. a comparable phase of basin restriction in east greenland, albeit somewhat later (latest toarcian – aalenian), appears to be indicated by evidence of brackish water conditions in the lower part of the offshore marine sortehat formation (dam & surlyk 1998; koppelhus & hansen 2003, this volume). this event may be attributable to progressive tectonic isolation of the jameson land basin due to regional uplift farther north. basin evolution in both east greenland and denmark was thus highly similar with regional subsidence by thermal contraction following rift events in the latest palaeozoic and earliest mesozoic. a marked transgressive trend characterised deposition in both areas; marine conditions were restricted to the axial parts of the seaways in the late triassic and earliest jurassic, spreading to the basin margins (e.g. skåne, bornholm) and the most proximal depocentres (e.g. jameson land basin) by the latest sinemurian – early pliensbachian. stepwise regression and basin restriction in the danish basin in the middle–late toarcian probably resulted from progressive uplift of the ringkøbing–fyn high, heralding the regional mid-jurassic uplift event (nielsen 2003, this volume). mid-jurassic regional uplift and erosion in east greenland, the rhaetian – lower bajocian prerift megasequence is restricted to the jameson land basin, which contains a relatively complete succession. the youngest strata beneath the unconformity that caps the megasequence are of early bajocian age based on dinocysts (underhill & partington 1994; koppelhus & hansen 2003, this volume) and supported by sr-isotope data (m. engkilde, personal communication 2000). a detailed ammonite zonation has been established for the shallow marine pelion formation overlying the unconformity but the ammonites are strictly boreal and the bajocian–bathonian interval cannot be directly correlated with european zonations. however, the immediate predecessor of the oldest of the boreal middle jurassic ammonites, cranocephalites borealis, is the subgenus defonticeras of the genus sphaeroceras from the north-eastern pacific which is confidently dated to the uppermost lower bajocian stephanoceras humphriesianum chronozone (callomon 1985). the great resemblance of sphaeroceras (defonticeras) oblatum and c. borealis suggests that the age difference 12 between them is small and the age of the c. borealis zone and of the basal onlapping strata is thus early late bajocian. there is a short hiatus between the sortehat and pelion formations and combined evidence from dating by sr-isotopes, dinocysts and ammonites suggests that it covers an interval across the lower–upper bajocian boundary. the hiatus is associated with a complete change in basin configuration and drainage pattern marking the onset of rifting in east greenland. the overlying deposits of the pelion formation and its correlatives show extensive onlap onto basin margins and northwards up the axis of the new rift (alsgaard et al. 2003, this volume; engkilde & surlyk 2003, this volume; larsen et al. 2003, this volume). the base of this early syn-rift succession youngs to the north and onlaps progressively older rocks from upper triassic through lower triassic and upper permian to crystalline basement in a northwards direction (surlyk 2003, this volume). the absence of lower jurassic rocks north of jameson land has been interpreted to reflect large-scale, possibly domal, uplift of northern east greenland in late early jurassic time (surlyk 1977a, b; surlyk et al. 1993). a similar situation is known from the norwegian side of the rift complex where large areas were uplifted in late early jurassic time; the stratigraphy on andøy on the conjugate margin of northern east greenland thus shows the same development as in wollaston forland, i.e. crystalline basement draped by a thin veneer of upper palaeozoic carbonates is directly overlain by middle jurassic sandstones (dalland 1981). the reality of early jurassic uplift to the north of jameson land originally suggested on stratigraphic grounds has recently been corroborated on the basis of apatite fission track thermochronology by johnson & gallagher (2000). in contrast, the jameson land area shows no evidence of early jurassic uplift and cooling (mathiesen et al. 2000). data from the norwegian shelf show that wide areas were uplifted in late early jurassic time (doré 1992). it has long been known that major uplift took place in the north sea in late early – early middle jurassic times (whiteman et al. 1975; hallam & sellwood 1976; eynon 1981; ziegler 1988; underhill & partington 1993, 1994). the uplifted area is generally referred to as the ‘midnorth sea dome’ and has been interpreted as having been caused by pre-rift heating and uplift followed by volcanism and rifting. the uplifted area underwent strong erosion and gradually deflated, being onlapped and subsequently flooded during middle and late jurassic times. underhill & partington (1993, 1994) demonstrated that the strata subcropping the erosional unconformity became gradually older approaching the centre of the uplift while the onlapping strata became younger in the same direction. more recent work has shown that the uplifted area was not a simple well-defined dome but involved the ringkøbing–fyn high, much of the danish basin and the fennoscandian border zone (nielsen 1995, 2003, this volume; andsbjerg et al. 2001). in the north sea, the main unconformity is typically constrained to the mid-aalenian in marginal areas. in the sorgenfrei–tornquist zone of the danish basin, for example, lower aalenian strata both underlie and overlie the unconformity and, based on dinoflagellate data, the unconformity appears to lie within the upper levels of the lowermost aalenian l. opalinum chronozone (nielsen 2003, this volume). as in east greenland, however, the limitations of the biostratigraphic data should be acknowledged since dating of the dominantly terrestrial sediments of the strata overlying the unconformity in the north sea is notoriously difficult. although the uplift and especially the onset of rifting and the associated radical basin reorganisation appear to have occurred earlier in the danish region than in east greenland, the main onlap phase onto the regional unconformity was broadly coeval, from the late bajocian to the early oxfordian (underhill & partington 1993, 1994; andsbjerg et al. 2001; andsbjerg & dybkjær 2003, this volume; nielsen 2003, this volume; surlyk 2003, this volume). it is noteworthy that this northwards delay in initial uplift from the danish region to east greenland is mirrored passing southwards from the danish area to the dutch sector of the central graben where the uplift – rift onset hiatus spans the mid-bajocian – mid-callovian (herngreen et al. 2003, this volume). regional uplift, erosion and subsequent subsidence, onset of rifting and onlap of the previously uplifted area thus took place in both east greenland and denmark in late early – early middle jurassic times. a common cause can be envisaged for both regions but the succession of events seems to be delayed in east greenland compared to the central north sea. syn-rift megasequence the regional uplift event at the early–middle jurassic transition was succeeded by the onset of a long-term rifting episode, which began in the middle jurassic, peaked in the late jurassic and persisted into the earliest cretaceous. rifting was not continuous but comprised phases of more intense rifting and block rotation 13 alternating with more tranquil periods of regional subsidence. the syn-rift succession can be subdivided into a number of tectonostratigraphic units marked by rift events followed by more gradual subsidence. seven regional tectonostratigraphic sequences have been recognised for the aalenian–valanginian stratigraphic interval in the central and northern north sea (uk and norwegian sectors) by rattey & hayward (1993); the evolution of the danish central graben in the aalenian–ryazanian described by andsbjerg & dybkjær (2003, this volume) is broadly compatible with the regional framework of rattey & hayward (1993) although megasequences were not defined in the danish central graben study. the seven north sea tectonostratigraphic sequences of rattey & hayward (1993) are equivalent to six sequences for the correlative interval in east greenland (surlyk & noe-nygaard 2000; surlyk 2003, this volume). the main difference between the two regions seems to be the delayed onset and culmination of rifting in east greenland compared to the north sea and the apparent lack of a tectonostratigraphic sequence boundary roughly at the oxfordian–kimmeridgian boundary in east greenland. otherwise the sequences correspond broadly to each other in timing and stratigraphic development, and it is not always clear if the differences, i.e. the slightly older positions of the boundaries in the north sea, are real. they may also reflect dating by ammonites in greenland and by dinocysts in the north sea, respectively, and uncertainties in the correlation between the two zonations. it has been noted by a few workers that the base of dinocyst zones tend to occur at progressively higher levels compared to ammonite zone boundaries in a north sea – east greenland – north greenland tran14 ryazanian volgian berriasian tithonian kimmeridgian oxfordian callovian bathonian bajocian aalenian toarcian pliensbachian sinemurian hettangian rhaetian c re ta ce ou s ju ra ss ic tr ia ss ic lo w er lo w er u pp er m id dl e u pp er u m l u l u m l u m l u m l u l u m l u m l u l u l 210 200 190 180 170 160 150 140 chronostratigraphyma nesw danish central graben rfhrfh stz skp sw nedanish basin sy nri ft m eg as eq ue nc e pr eri ft m eg as eq ue nc e sy nri ft m eg as eq ue nc e pr eri ft m eg as eq ue nc e sect (smelror 1993; s. piasecki, personal communication 2002). the northwards younging of tectonostratigraphic boundaries may thus be real, possibly reflecting progressive northward propagation of the rift system as suggested by surlyk & clemmensen (1983), or apparent, reflecting correlation problems at a time of marked ammonite provinciality or northward migration of indicator dinocysts with respect to the more finely-tuned ammonite zonation that forms the basis for jurassic chronostratigraphy. early syn-rift sedimentation (late aalenian – bajocian) in the danish basin and the danish central graben was confined to narrow subsiding grabens and the succession is probably incomplete with a number of inferred unconformities within the bajocian–bathonian part. it is noteworthy that a prominent unconformity is recorded in the uppermost bathonian of the northern danish central graben, recording a marked shift in subsidence patterns during early rifting (andsbjerg 2003, this volume). although loosely constrained biostratigraphically, a hiatal surface is also inferred at this level in the danish basin (fig. 2; nielsen 2003, this volume). this event is not detected in the east greenland sedimentary record where the bathonian–callovian transition is characterised by transgression and progressive backstepping of sedimentary systems (fig. 2); minor hiatuses are recorded at this stratigraphic level but these resulted from condensation and non-deposition in offshore settings. the syn-rift successions of denmark and east greenland are suggestive of a northward younging diachroneity of rift phases. the rift climax occurred in the early oxfordian – middle middle volgian in the danish sector of the central graben, albeit with an important lull in the late kimmeridgian characterised 15 no data no data s n source rock estuarine/lagoonal sandstones, heteroliths, mudstones and coal beds hiatus/condensed jameson land wollaston forland – kuhn øs n w e lacustrine deltas, sand-dominated lacustrine mudstones alluvial/delta plain – paralic, sand-dominated fluvial and estuarine sandstones, conglomerates floodplain mudstones shallow marine sandstones offshore/basinal mudstones, heteroliths organic-rich offshore/basinal mudstones deep marine sandstones deep marine conglomerates coal sy nri ft m eg as eq ue nc e pr eri ft m eg as eq ue nc e sy nri ft m eg as eq ue nc e fig. 2. chronostratigraphic scheme of the uppermost triassic – lowermost cretaceous of the danish central graben, the danish basin and east greenland showing the main lithologies, depositional environments and tectonostratigraphic sequences. simplified from andsbjerg & dybkjær (2003, this volume), nielsen (2003, this volume) and surlyk (2003, this volume); time-scale after gradstein et al. (1994). rfh, ringkøbing–fyn high; skp, skagerrak–kattegat platform; stz, sorgenfrei–tornquist zone. by regression and shoreface progradation (andsbjerg & dybkjær 2003, this volume; johannessen 2003, this volume; møller & rasmussen 2003, this volume). the late middle and late volgian saw a general waning in rift activity in the danish central graben resulting in the development of more symmetrical sub-basins, associated with a general reduction in the overall sedimentation rate. indeed, the upper volgian – lower ryazanian in the central graben is characterised by a relatively condensed stratigraphic package of organic-rich ‘hot shales’, associated locally with basin floor sands (donovan et al. 1993; ineson et al. 2003, this volume) an oxfordian – early volgian rift climax seems to be applicable to the jameson land basin at the southern end of the east greenland rift basin where chaotic deepwater sandstones of the upper oxfordian – lower volgian hareelv formation mark the rift climax (surlyk & noe-nygaard 2001; surlyk 2003, this volume). it was succeeded by rapid progradation and basin infill in middle and late volgian times as rift activity waned. in wollaston forland at the northern end of the rift basin, however, the late oxfordian – early volgian was characterised by gentle block tilting, whereas the rift climax accompanied by strong block tilting took place in the middle volgian. taken at face value, therefore, the stratigraphic synrift histories of the two regions are broadly similar but the main events appear to have started earlier in the south. conclusions unravelling the complexities of the jurassic rifted seaway in the north atlantic region continues to be a subject of major research interest, not least due to the hydrocarbon potential of jurassic basins on both sides of the atlantic ocean. the basins of east greenland and denmark represent important pieces in this jigsaw puzzle and the studies reported in the following papers will help to further constrain regional models of rift development, and to better understand jurassic stratigraphic development in general. comparison of the jurassic evolution of these areas makes it possible to construct an integrated picture of the long, relatively narrow seaways, allowing testing of ideas concerning basin evolution, domal versus regional uplift, and timing of the onset and climax of rifting. in addition to this regional perspective, parallel research into the jurassic of the east greenland and danish basins allows comparison of the signatures of sequence stratigraphic surfaces as seen on well logs, in cores and at outcrop, and of sequences recognised and defined on the basis of very different data types. furthermore, experience gleaned from the extensive outcrops of east greenland aids interpretation of restricted outcrops on bornholm and in skåne and allows them to be placed within a regional framework. the tectonostratigraphic summary presented above shows that the main tectonic events and stratigraphic trends in east greenland, denmark and adjacent areas are highly similar but apparently somewhat out of phase for the syn-rift successions (fig. 3). the rhaetian – early jurassic was characterised by regional subsidence following late palaeozoic and early mesozoic rift events and the detailed stratigraphic signature reflects relative sealevel changes superimposed on a long-term sea-level rise. major regional uplift heralding the onset of rifting took place broadly at the early–middle jurassic boundary and the uplifted areas underwent marked erosion. subsequent subsidence began in the aalenian in the north sea and in the late bajocian in east greenland concomitant with the onset of rifting, resulting in major regional onlap and transgression. continued relative sea-level rise, reflecting the early rifting, took place in the bathonian; deltas and shallow marine sandy systems were drowned almost everywhere by the end of the callovian. rifting culminated in early oxfordian – volgian times with major block tilting and deposition of fault-scarp aprons and basin-floor fans. the rift climax was delayed in northern east greenland compared to areas further south. the timing and style of tectonic events thus exerted the main control on the long-term trends in stratigraphic evolution. in the early jurassic, however, relative sealevel changes that were unrelated to local tectonics seem to have exerted the main control on the depositional motifs (dam & surlyk 1995, 1998; andsbjerg & dybkjær 2003, this volume; nielsen 2003, this volume; surlyk 2003, this volume). the late jurassic deepening and transgressive trend, on the other hand, appears to reflect accelerated regional subsidence, increased tilting of fault blocks, eustatic sea-level rise or a combination of these factors, and isolation of the dominant control is difficult without comparison with successions on other lithospheric plates. acknowledgements this introductory paper is based mainly on the detailed studies reported in this book; we acknowledge the authors and thank peter r. dawes, peter n. johannessen, michael larsen and lars h. nielsen for useful comments. 16 17 lo ca lis ed fa ul t re -a ct iv at io n r eg io na l th er m al su bs id en ce u pl ift a nd de ep e ro si on fa ul tco nt ro lle d su bs id en ce o f ea st er n d c g m in or u pl ift (s øg ne b as in ) re gi on al ba ck st ep pi ng m aj or h al fgr ab en de ve lo pm en t ri ft hi at us , s ed im en ta ry pr og ra da tio n sh el f d ro w ni ng sh el f d ro w ni ng sh el f d ro w ni ng sh el f d ro w ni ng r en ew ed r ift s ub si de nc e – m ar ke d ba si n se gm en ta tio n an d bl oc k ro ta tio n pe ak tr an sg re ss io n u pl ift o f r fh , n etil tin g of b as in , de ep e ro si on t o sw fa ul tco nt ro lle d su bs id en ce o f s t z re gi on al ba ck st ep pi ng r en ew ed r eg io na l su bs id en ce , in cl ud in g r fh m ar in e in un da tio n o ve ra ll tr an sg re ss ive tr en d m in or u pl ift , ba si n re or ga ni sa tio n g en tle b lo ck r ot at io n, in iti al s ed im en ta ry pr og ra da tio n re gi on al ba ck st ep pi ng d ee pe ni ng – p er si st en t re gi on al s ub si de nc e (b lo ck r ot at io n) pr og ra da tio n of sh el f s ys te m s u pl ift a nd in ci si on in fil l/d ra pe o f er os io na l r el ie f ba si n se gm en ta tio n an d bl oc k ro ta tio n u pl ift /e ro si ono nl ap o nt o pr eju ra ss ic ba se m en t, se di m en ta ry pr og ra da tio n g en tle b lo ck ro ta tio n re gi on al ba ck st ep pi ng sy nri ft p ha se pr eri ft p ha se u pl ift fa ul tco nt ro lle d su bs id en ce ( m in or /m aj or ) d en m ar k ea st g re en la nd ja m es on l an d – m iln e la nd w ol la st on f or la nd – k uh n ø c en tr al g ra be n d an is h ba si n r ift cl im ax ?u pl ift r ift cl im ax r ift cl im ax r eg io na l th er m al su bs id en ce r eg io na l th er m al su bs id en ce c hr on os tr at ig ra ph y m a r ya za nia n vo lg ia n be rr ias ia n t ith on ia n k im m er id gi an o xf or di an c al lo vi an ba th on ia n ba jo ci an a al en ia n to ar ci an pl ie ns ba ch ia n si ne m ur ia n h et ta ng ia n r ha et ia n cretaceous jurassic triassic lower lowerupper middle upper u m l u l u m l u m l u m l u l u m l u lm u l u l 21 0 20 0 19 0 18 0 17 0 16 0 15 0 14 0 fi g. 3 . sc h em e sh o w in g th e m ai n t ec to n ic e ve n ts a n d t re n d s in b as in e vo lu tio n i n t h e d an is h c en tr al g ra b en , th e d an is h b as in a n d e as t g re en la n d p lo tt ed o n a t im e ax is . t im esc al e af te r g ra d st ei n e t a l. 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(eds): siliciclastic sequence stratigraphy: recent developments and applications. american association of petroleum geologists memoir 58, 449–484. whiteman, a.j., rees, g., naylor, d. & pegrum, r.m. 1975: north sea troughs and plate tectonics. norges geologiske undersøkelse 316, 137–161. ziegler, p.a. 1988: evolution of the arctic – north atlantic and the western tethys. american association of petroleum geologists memoir 43, 198 pp. ziegler, p.a. 1990: geological atlas of western and central europe, 2nd edition, 239 pp. amsterdam: elsevier for shell internationale petroleum maatschappij. geological survey of denmark and greenland bulletin 1, 367-402 367 sedimentology and sequence stratigraphy of paralic and shallow marine upper jurassic sandstones in the northern danish central graben peter n. johannessen paralic and shallow marine sandstones were deposited in the danish central graben during late jurassic rifting when half-grabens were developed and the overall eustatic sea level rose. during the kimmeridgian, an extensive plateau area consisting of the heno plateau and the gertrud plateau was situated between two highs, the mandal high to the north, and the combined inge and mads highs to the west. these highs were land areas situated on either side of the plateaus and supplied sand to the gertrud and heno plateaus. two graben areas, the feda and tail end grabens, flanked the plateau area to the west and east, respectively. the regressive–transgressive succession consists of intensely bioturbated shoreface sandstones, 25–75 m thick. two widespread unconformities (sb1, sb2) are recognised on the plateaus, forming the base of sequence 1 and sequence 2, respectively. these unconformities were created by a fall in relative sea level during which rivers may have eroded older shoreface sands and transported sediment across the heno and gertrud plateaus, resulting in the accumulation of shoreface sandstones farther out in the feda and tail end grabens, on the south-east heno plateau and in the salt dome province. during subsequent transgression, fluvial sediments were reworked by high-energy shoreface processes on the heno and gertrud plateaus, leaving only a lag of granules and pebbles on the marine transgressive surfaces of erosion (mtse1, mtse2). the sequence boundary sb1 can be traced to the south-east heno plateau and the salt dome province, where it is marked by sharp-based shoreface sandstones. during low sea level, erosion occurred in the southern part of the feda graben, which formed part of the gertrud and heno plateaus, and sedimentation occurred in the norwegian part of the feda graben farther to the north. during subsequent transgression, the southern part of the feda graben began to subside, and a succession of backstepping back-barrier and shoreface sediments, 90 m thick, was deposited. in the deep tail end and feda grabens and the salt dome province, sequence boundary sb2 is developed as a conformity, indicating that there was not a significant fall in relative sea level in these grabens, probably as a result of high subsidence rates. backstepping lower shoreface sandstones overlie sb2 and show a gradual fining-upwards to offshore claystones that are referred to the farsund formation. on the plateaus, backstepping shoreface sandstones of sequence 2 are abruptly overlain by offshore claystones, indicating a sudden deepening and associated cessation of sand supply, probably caused by drowning of the sediment source areas on the mandal, inge and mads highs. during the volgian, the gertrud plateau began to subside and became a graben. during the late kimmeridgian – ryazanian, a long-term relative sea-level rise resulted in deposition of a thick succession of offshore claystones forming highstand and transgressive systems tracts on the heno plateau, and in the gertrud, feda and tail end grabens. keywords: north sea, danish central graben, kimmeridgian, sedimentology, sequence stratigraphy, ichnology, back-barrier – shoreface sediments, palaeogeography, sand distribution geological survey of denmark and greenland, geocenter copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: pjo@geus.dk geological survey of denmark and greenland bulletin 1, 367–402 (2003) © geus, 2003 368 kimmeridgian shoreface sandstones form thick (25–200 m), widespread units in the north sea central graben and constitute good reservoirs in the uk sector (fulmar formation; fife, clyde, and fulmar fields) and in the norwegian sector (ula formation; ula and gyda fields). in the danish sector of the central graben, two fields are found: the gert/mjølner oil field and the elly gas field, which are expected to start producing in the near future. in order to locate additional hydrocarbon fields with similar reservoir sandstones, it is of major importance to understand the distribution of the sandstones and their architecture. the reservoir sandstones in most of the study area are deeply buried (3.5–5 km below mean sea level (b. msl)), resulting in limited seismic resolution. furthermore, a large number of closely-spaced faults in the area render it difficult to trace seismic reflectors. in general, the biostratigraphic resolution is too poor to resolve correlations within the kimmeridgian sandy successions; a sequence stratigraphic study based on cores and gert ridge 4°e 56°n mandal high mid north sea high tail end g raben ringkøbing–fyn high salt dome province søgne basin gertrud graben/ plateau feda graben heno plateau m ads h igh arne–elin graben outer rough basin inge high ål basin 50 km 56°n ravn-2ravn-2 2/11-7 a a´ b´ 4°e elly-2 diamant-1 nora-1 w-1 gwen-2 gert-1 eg-1 skarv-1 gert-4 jeppe-1 falk-1 ravn-1 2/12-1 q-1 ravn-2 b karl-1 50 km gert-2 edna-1 elly-1 ■ ■ ■■ ■ ■ ■■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■■ ■■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ 55°n 4°e 100 km salt structures normal fault reverse fault well structural high outer moray firth v ik in g g ra be n central graben ringkøbing– fyn high n g nl uk dk mid north sea high fig. 1. maps showing the regional location and structural framework of the danish central graben and the locations of wells and geosections (a–a′, b–b′; fig. 2) in the northern part of the graben. national sectors of the north sea: dk, denmark; g, germany; n, norway; nl, the netherlands; uk, united kingdom. 369 petrophysical logs has thus been undertaken. all available cores were described and interpreted sedimentologically, including a detailed study of trace fossils and ichnofabric (johannessen 1995). petrophysical logs and cuttings samples were interpreted by comparison with the cored sections. north sea boreal age/stage terminology (sensu gallico) has been used throughout the paper. the aim of this paper is to describe and interpret the depositional environments of the kimmeridgian sediments on the gertrud and heno plateaus, in the tail end and feda grabens and in the salt dome province. the results of a detailed sequence stratigraphic analysis are presented as a series of log-panels and palaeogeographic maps and the potential distribution of further reservoir sandstones is discussed. geological setting the danish part of the north sea central graben is bordered by the mid north sea high to the west and gert-1 gert-4 jeppe-1 karl-1 mid north sea high feda graben inge high ål basin gertrud graben mandal high karl platform søgne basin coffee soil fault ringkøbing–fyn high (east north sea high) ravn-1 ravn-2 w-1 nora-1 a a´ heno plateau tail end graben ringkøbing–fyn high arne–elin graben mads high coffee soil fault post-ryazanian lower volgian – ryazanian callovian – upper kimmeridgian aalenian–callovian triassic – lower jurassic zechstein pre-zechstein 1 0 2 3 4 5 6 7 8 9 10 20 30 40 d ep th ( km ) 50 60 70 (km) wsw ene 1 0 2 3 4 5 6 7 8 9 10 20 30 d ep th ( km ) 40 50 (km) wsw ene b b´ gert ridge fig. 2. geosections oriented perpendicular to the axis of the danish central graben. for location, see fig. 1. the locations of wells used in this study that lie on or adjacent to the geosections are indicated; the well traces shown with solid lines are on the section line, those shown by dashed lines are projected into the line of section. modified from møller & rasmussen (2003, this volume). the ringkøbing–fyn high to the east (fig. 1). the central part of the area is dominated by the deep tail end graben, which is the main jurassic depocentre with more than 4 km of primarily upper jurassic sediments (fig. 2; jensen et al. 1986; damtoft et al. 1992; johannessen & andsbjerg 1993; andsbjerg & dybkjær 2003, this volume; japsen et al. 2003, this volume; møller & rasmussen 2003, this volume). the middle jurassic succession thickens eastwards towards the main fault against the ringkøbing–fyn high, suggesting that middle jurassic sediments were deposited during the early phase of half-graben subsidence (damtoft et al. 1992; johannessen & andsbjerg 1993; andsbjerg 2003, this volume; andsbjerg & dybkjær 2003, this volume). rift-related subsidence within the danish portion of the central graben is believed to have started in the bajocian and the half-graben continued to subside until the ryazanian (fig. 3). during the bajocian–volgian, four major half-graben systems were formed and each new half-graben propagated farther westwards (fig. 2; damtoft et al. 1992). from east to west, these half-graben systems are termed the søgne basin – tail end graben – salt dome province trend, the gertrud graben, and the ål basin – outer rough basin trend (fig. 1). the feda graben is an exception to this westward propagation, in that its northern part probably started to subside during middle jurassic times (damtoft et al. 1992; rasmussen 1995). it has been proposed earlier that the overall jurassic transgression proceeded from south to north through the central graben towards the triple junction dome (eynon 1981; koch 1983). more recently, however, on the basis of more extensive core data and better biostratigraphic data, it has been suggested that the transgression came from the north (johannessen & andsbjerg 1993). during late jurassic time, after transgression of the deepest part of the incipient half-graben system, the rest of the danish central graben was transgressed from the east towards the west (fig. 3), with the exception of certain of the intrabasinal highs – the mandal, inge and mads highs (damtoft et al. 1992; johannessen & andsbjerg 1993). the heno plateau is a complex block-faulted platform area with a much thinner jurassic succession than in the graben areas, generally less than 1000 m thick (fig. 2). water depths during the late kimmeridgian were particularly shallow in the northern and central parts of the heno plateau (diamant-1, eg-1, ravn-1, -2, w-1). farther to the south-east (elly-1, -2, falk-1, skarv-1), the heno plateau experienced a higher degree of subsidence resulting in increased accommodation space and greater water depths. thus, although structurally forming part of the plateau area, the depositional environments tended towards those of the deeper graben areas such as the tail end graben and the salt dome province lying 370 system stage lo w er ju ra ss ic u pp er m i d d l e åsgard formation leek member bo member heno fm lola formation bryne formation offshore organic-rich marine shales marine mudstones and siltstones submarine fan sandstones and siltstones shallow marine sandstones and siltstones, subordinate back-barrier sediments paralic and non-marine sandstones, siltstones, mudstones and coals marine calcareous mudstones and marlstones lithostratigraphy ryazanian volgian kimmeridgian oxfordian callovian bathonian bajocian aalenian c re ta ce ou s valanginian vyl fm poul fmfarsund formation lulu formation u l u l l u u m l l m u m l u l u u l u m l hiatus fig. 3. middle jurassic – lowermost cretaceous stratigraphic scheme for the danish central graben. modified from michelsen et al. (2003, this volume). east and south-east of the heno plateau (figs 1, 2). on the mads high, upper cretaceous chalk unconformably overlies palaeozoic rocks (damtoft et al. 1992). the inge and mads highs formed a continuation of the mid north sea high during bajocian–kimmeridgian times, and constituted a large positive area that possibly was an important sediment source area (johannessen et al. 1996). during the oxfordian and kimmeridgian, the gertrud plateau and the heno plateau together formed a flatlying, slowly subsiding platform (fig. 1; söderström et al. 1991; rasmussen 1995; johannessen et al. 1996). the gertrud graben (previously the site of the gertrud plateau) did not begin to subside significantly before the volgian. the mandal high, which straddles the norwegian– danish border east of the gertrud plateau and becomes pronounced in the norwegian sector farther north, may have been an important source area for the upper jurassic sandstones (figs 1, 2). during the late kimmeridgian and volgian, a longterm relative sea-level rise resulted in deposition of thick offshore claystones in the central part of the central graben (heno plateau and gertrud and feda grabens; damtoft et al. 1992; johannessen & andsbjerg 1993; johannessen et al. 1996). during the volgian and early cretaceous, the transgression continued towards the west, and shallow marine sandstones may have been deposited in the outer rough basin and ål basin at the western margin of the central graben fringing the mid north sea high (figs 1–3), as seen in the adjacent uk sector (‘fife sandstones’; mackertich 1996; spathopoulos et al. 2000). regional stratigraphy the upper jurassic succession of the danish central graben consists of the lola, heno, farsund and poul formations (figs 3, 4; vollset & doré 1984; jensen et al. 1986; michelsen et al. 2003, this volume). the lola formation occurs in the southern part of the gertrud and heno plateaus and in the tail end graben and is characterised by offshore claystones. it is mainly of oxfordian age but extends up into the upper kimmeridgian in the feda graben (see below). the heno formation is kimmeridgian in age and is characterised by shoreface sandstones showing an overall regressive to transgressive pattern on the heno and gertrud plateaus. in the deeper parts of the south-eastern heno plateau, the feda graben and the salt dome province, the shoreface sandstones become more finegrained, consisting of overall regressive to transgressive clayey shoreface sandstones. shoreface sandstones are not seen in wells in the tail end graben. the ‘basal sandstone unit’, and the lola and heno formations have been described previously from the feda graben and gertrud plateau area (bergan et al. 1989; söderström et al. 1991; johannessen et al. 1996). in the danish sector, the back-barrier and shoreface sediments previously referred to the ‘basal sandstone unit’ are redefined in this volume as the gert member of the heno formation (fig. 4); the shoreface sandstones classically referred to the heno formation (jensen et al. 1986; johannessen et al. 1996) are referred to the ravn member of the heno formation (michelsen et al. 2003, this volume). new biostratigraphical data show that the lola formation on the heno plateau, the southern part of the gertrud plateau and in the tail end graben is not younger than early kimmeridgian (cymodoce chronozone; andsbjerg & dybkjær 2003, this volume). however in the gert area, where lola formation mudstones occur sandwiched between the gert and ravn members of the heno formation (fig. 4), the succession is no older than late kimmeridgian (mutabilis chronozone; johannessen et al. 1996). consequently, the offshore claystones referred to the lola formation in the gert area are not age-equivalent with the lola formation on the heno plateau, as previously suggested (johannessen & andsbjerg 1993). the farsund formation is of late kimmeridgian – early ryazanian age and consists of offshore claystones that are widespread in the danish central graben, although absent over certain intra-basinal highs. the poul formation (volgian) is recognised in the eastern part of the tail end graben, where it consists of turbidite sandstones (jensen et al. 1986; damtoft et al. 1992); this unit is not described further in this paper. regionally in the central graben, the lola formation correlates with the haugesund formation (norway) and the heather formation (uk sector; jensen et al. 1986). the heno formation is roughly equivalent to the fulmar formation in the uk sector (johnson et al. 1986; armstrong et al. 1987; donovan et al. 1993; partington et al. 1993; price et al. 1993; wakefield et al. 1993) and shows similarities with the ‘heno equivalent’ and the ula formation of the norwegian sector (bergan et al. 1989; taylor & gawthorpe 1993; howell et al. 1996). the farsund formation correlates with the kimmeridge clay formation of the uk sector and the farsund formation of the norwegian sector (vollset & doré 1984; jensen et al. 1986). 371 372 fa rs un d fo rm at io n 200 m claystone clayey sandstone sandstone danish central graben 20 km gertrud plateau/graben gr dt res gert-1 gr dt res jeppe-1 gr dt res gwen-2 t op farsund fm gert-4 gr dt res elly-2 gwen-2 gert-1 gert-4 jeppe-1 edna-1 ravn-1 h en o fm h en o fm r av n m b g er t m b lo la f m feda graben fig. 4. correlation of the upper jurassic – lowermost cretaceous formations in the northern part of the danish central graben, illustrated by gamma-ray (gr), sonic velocity (dt) and resistivity (res) logs. the gert and ravn members of the heno formation are defined in michelsen et al. (2003, this volume). b. fm, bryne formation. 373 fa rs un d fm h en o fm (r av n m b) lo la f m b. fm? fa rs un d fo rm at io n lo la f or m at io n fj er ri ts le v fo rm at io n heno plateau salt dome province gr dt res edna-1 ravn-1 gr dt res elly-2 gr dt res 374 4840 m b. msl mtse sb0+sb1 5316 m b. msl 4938 m b. msl fs fs fs mtse sb 2+mtse 2 fs 4644 m b. msl td log 4922 m b. msl 2/11-7 danish central graben gwen-2 gert-1 gert-2 ravn-1w-1 gert-4 2/12-1 jeppe-1 feda graben q-1 gert-2 gr gert-4 gr gert-1 gr 2/12-1 gr 2/11-7 gr 100 m mfs1 fs sb2+ mtse2 tst hst tst tst tst sb0+sb1 mtse 20 km fig. 5. gamma-ray log panel extending from the feda graben to the gertrud and heno plateaus showing the upper jurassic sand-rich succession in the northern part of the danish central graben. the locations of cored sections are shown; the accompanying legend also applies to figs 8, 13 and 17. note that the flooding surface used as the datum represents the upper limit of detailed investigation in this study; andsbjerg & dybkjær (2003, this volume) interpret this surface as a composite flooding surface and sequence boundary in the gwen-2 and ravn-1 wells. 375 heno plateau 5002 m b. msl 4319 m b. msl 4201 m b. msl 4205 m b. msl 4266 m b. msl sb 1 sb2+mtse2 sb 0 mfs1 tst hst tst hst lst sedimentary environments back-barrier sediments shoreface conglomerate middle–upper shoreface sandstone lower shoreface clayey sandstone offshore claystone systems tracts lst lowstand systems tract tst transgressive systems tract hst highstand systems tract key surfaces sb0 sequence boundary, base jurassic unconformity mtse marine transgressive surface of erosion sb1 sequence boundary of sequence 1 mfs1 maximum flooding surface in sequence 1 fs flooding surface onlap truncation additional features core casing point 4840 m b. msl base jurassic, metres below mean sea level normal fault total depth mtse 1 gertrud plateau ravn-1 gr w-1 gr q-1 gr gwen-2 gr jeppe-1 gr fs sb 2+mtse2 sb1+mtse1 mfs1 mfs0 sb0 td facies and depositional environments three broad lithofacies associations have been recognised: offshore claystones, back-barrier sediments and shoreface sandstones (see johannessen 1995 for detailed facies descriptions). these are described and interpreted briefly here in terms of four recurrent elements that form the building blocks of the sequence stratigraphic framework: (1) offshore claystones, (2) transgressive back-barrier to shoreface deposits, (3) regressive–transgressive shoreface sandstones and conglomerates and (4) sharp-based lower shoreface sandstones. offshore claystones description the offshore claystones are characterised by high gamma-ray values, high resistivity readings and low sonic velocities (fig. 4). cuttings samples show that the claystones are associated with thin dolomite, sandstone and siltstone beds, and also that the organic content is variable. the claystones of the lola formation contain larger amounts of terrestrially-derived material than those of the farsund formation (jensen et al. 1986; michelsen et al. 1987). the claystones contain abundant and high diversity dinoflagellate assemblages. interpretation the abundant dinoflagellates indicate that the claystones were deposited in a fully marine offshore setting. the claystones of the lola formation were probably deposited nearer to the sediment source area than those of the farsund formation, indicating an overall transgressive trend (jensen et al. 1986; michelsen et al. 1987). transgressive back-barrier deposits to shoreface sandstones description the lowermost 56 m of the upper jurassic section were cored in gert-1, situated in the eastern part of the feda graben near the gertrud plateau (figs 5–7). the sec376 mtse sb mfs conglomerate outsized quartz clast sandstone silt–claystone heterolith coal coal clast pyrite cross-bedding low-angle cross-bedding cross-lamination climbing ripple cross-lamination faint parallel lamination water escape structure lithology sedimentary structures asterosoma isp. chondrites isp. diplocraterion isp. helminthopsis isp. ophiomorpha nodusa palaeophycus heberti planolites isp. rhizocorallium isp. skolithos isp. teichichnus isp. terebellina isp. thalassinoides isp. escape trace rootlets bivalve shell leaf biogenic structures/components additional features a fig. core photos blow up of sedimentological log marine transgressive surface of erosion sequence boundary maximum flooding surface transgressive systems tract highstand systems tract iii intense ii moderate bioturbation i weak tst hst facing page: fig. 6. sedimentological core log from sequence 1 in the gert-1 well, located in the feda graben near the gertrud plateau. typical small-scale coarsening-upwards to fining-upwards successions recognised in the cored section are illustrated in the enlarged logs (a, b). the accompanying legend (above) also applies to figs 9 and 11. 377 40 30 20 10 0 t st ba ck -b ar ri er sb1/sb0 4940 m core depth 4930 m core depth sand si pblcl sand si pblcl 0 2 4 m m 0 jurassic permian passive channel fill active channel fill mouth bar passive channel fill active channel fill mouth bar sand si pblcl gert-1 75 150 m a b a 4975 m core depth 0 gamma-ray 50 b fig. 7 378 core 5 box 10 gert-1 core 5 box 11 core 5 box 9 core 5 box 12 top base pla. rootlets claystone shells oph. pla. oph. crossbedded sandstone claystone pal. oph. oph. oph. 10 cm tion comprises interbedded fine-grained sandstones and claystones with coal beds, 2–15 cm thick. the fine-grained sandstones are typically so intensively bioturbated and affected by water-escape structures that all primary sedimentary structures are destroyed. numerous burrows of ophiomorpha nodosa as well as rootlets and coal fragments are recognised throughout the cored section (figs 6, 7). burrows of palaeophycus isp., diplocraterion isp., terebellina isp., skolithos isp. and planolites isp. also occur. the organic matter in the sandstone beds is solely composed of black, oxidised wood particles. the organic matter in the claystone beds and laminae is strongly dominated by terrestrially-derived particles, such as brown and black wood, cuticles, spores and pollen. marine dinoflagellate cysts are present but are scarce in most samples (0–3% of total palynomorphs); carboniferous spores and pollen occur abundantly (johannessen et al. 1996). coarseningto fining-upwards successions, 3–8 m thick, are recognised (figs 6, 7). the lower, coarseningupwards portion consists mainly of ophiomorpha-burrowed siltstones and fine-grained sandstones showing faint parallel bedding or low-angle cross-bedding, usually associated with water escape structures. the overlying fining-upwards portion is characterised by high-angle cross-bedded or low-angle cross-bedded fine-grained or medium-grained sandstones, overlain by structureless fine-grained sandstones, siltstones and claystones, in places capped by coal beds, 2–15 cm thick. the fining-upwards succession is usually associated with ophiomorpha nodosa burrows, water escape structures and rootlets. the lowermost 60 m thick section of gert-1 is characterised by a serrated gamma-ray log pattern (fig. 5). above this serrated interval is a unit (18 m thick) with consistently low gamma-ray readings, suggesting cleaner and probably more coarse-grained sandstones than below. a succession of very fine-grained sandstones, 6 m thick, overlies the clean, coarser-grained sandstones and is characterised by higher gamma-ray readings. the fine-grained sandstone unit is overlain by offshore claystones showing high gamma-ray values (fig. 5). interpretation a marginal marine setting is indicated by the association of rootlets and in situ coals, together with bioturbated sediments containing ophiomorpha nodosa. the numerous water escape structures indicate rapid deposition. the nature of the kerogen indicates a low-salinity, nearshore setting with a high input of terrestrial material and the small number of dinoflagellate cysts, and the dominance amongst these of a single genus (sentusidinium), indicates a low salinity, highly-stressed environment (noe-nygaard et al. 1987; smelror & leereveld 1989). the association of restricted marine environments indicates that the sediments were deposited within a back-barrier setting. the coarsening-upwards successions were probably deposited by prograding mouth bars, near the bayhead shoreline. the fining-upwards successions were deposited in channels by migrating mega-ripples or bars. the overlying fine-grained sandstones, siltstones and claystones represent the passive channel fill. the thin coal beds with rootlets, represent the final phase of abandonment. the occurrence of ophiomorpha nodosa in sandstones of the passive channel fill may indicate some marine influence. the channels are interpreted as distributary channels that dissected their previously deposited mouth bars. the general log motifs above the cored section in gert-1 suggest a cleaner sandstone unit abruptly overlying the back-barrier sediments indicating that it was deposited under high-energy conditions. further, the abundant dinoflagellate cysts in the overlying claystone suggest deposition in a fully marine offshore environment. hence, the clean sandstones are interpreted as backstepping shoreface sandstones deposited on a marine transgressive surface of erosion during a transgression (mtse; fig. 5). the occurrence of carboniferous spores and pollen suggests that carboniferous sediments may have been 379 facing page: fig. 7. core photo from sequence 1 of gert-1. bioturbated heterolithic sandstones overlie claystones and are in turn overlain by faintly parallel-laminated, fine-grained sandstones with burrows of ophiomorpha nodosa (oph.). cross-bedded fine-grained sandstones overlie the faint parallel-laminated sandstones. claystones with thick bivalve shells overlie the sandstones. the upper part of the sandstone unit, immediately below the claystones, has numerous rootlets. the coarseningupwards heterolithic sandstones and the faint parallellaminated sandstones represent a mouth bar deposited in a back-barrier setting, probably near a bayhead delta. the overlying cross-bedded sandstones were deposited in a distributary channel, cutting through its own mouth bar. the uppermost claystones with underlying rootlets represent the final abandonment of the channel. for position of core photo, see fig. 6. additional trace fossils: pal., palaeophycus heberti; pla., planolites isp. the dominant sediment source. thick carboniferous sandstones were cored below the upper jurassic sandstones in the nearby gert-2 well, situated on the upthrown side of the fault between gert-1 and gert-2, and may represent a local source of sandy sediment. further occurrences transgressive back-barrier deposits overlain by shoreface sandstones also occur in the lowermost part of the upper jurassic in the gert-2, jeppe-1, karl-1, 2/12-1, gert-4 and diamant-1 wells (figs 5, 8; söderström et al. 1991; johannessen & andsbjerg 1993; johannessen 1995; johannessen et al. 1996). the shoreface sandstones fine upwards in a stepwise manner to offshore claystones, suggesting continued transgression. on the gertrud plateau, the shoreface sandstones fine upwards into clayey sandstones deposited in a lower shoreface environment. the shoreface sandstones of diamant-1 on the westernmost part of the heno plateau are abruptly overlain by offshore claystones of the farsund formation 380 4205 m b. msl 3825 m b. msl 4362 m b. msl 4266 m b. msl 4365 m b. msl fs tst mtse1 mfs0 sb0+ sb1 * tst sb2+mtse2 sb 2 +mtse 2 sb1+mtse1 sb 0 100 m sb0+sb1 mfs 1 falk-1 elly-2 diamant-1 w-1eg-1 skarv-1 edna-1 ravn-1 ravn-2 ravn-1 gr w-1 gr eg-1 gr diamant-1 gr ravn-2 gr danish central graben 20 km heno plateau fig. 8. gamma-ray log panel of the upper jurassic sand-rich succession from the heno plateau to the salt dome province in the northern part of the danish central graben; for legend, see fig. 5. the falk-1, elly-2 and skarv-1 wells are located on the deeper, south-eastern part of the heno plateau. note that the flooding surface used as the datum represents the upper limit of detailed investigation in this study; andsbjerg & dybkjær (2003, this volume) interpret this surface as a composite flooding surface and sequence boundary in the falk-1, ravn-1 and ravn-2 wells. surface indicated with an asterisk is composite: mtse1 + sb2 + mtse2. (fig. 8). offshore claystones of the lola formation (fig. 8) were either never deposited in this area or were eroded during a subsequent fall in relative sea level. back-barrier sediments overlain by transgressive shoreface sandstones are not recorded on the remaining part of the heno plateau, the southern part of the gertrud plateau nor in the salt dome province and the tail end graben. regressive–transgressive shoreface sandstones and conglomerates description upper jurassic sandstones form a unit up to 70 m thick on the gertrud plateau, partly cored in the gert-2, jeppe-1 and gwen-2 wells (figs 5, 9, 10). the sandstones are very fineto medium-grained and form a coarsening-upwards succession; in the gwen-2 and jeppe-1 wells, this is abruptly overlain by a conglom381 4076 m b. msl 3987 m b. msl 3793 m b. msl 3855 m b. msl lstv mfs1 sb1 mfs0 sb0 fs fs sb2+ mtse2 lst salt dome province tst hst tst lst hst falk-1 gr elly-2 gr skarv-1 gr edna-1 gr erate unit, 0.5–3 m thick (figs 5, 9). the upper part of the sandstone succession, comprising mediumto finegrained sandstones, fines upwards over an interval of 15–30 m (fig. 5). the organic matter consists primarily of brown and black wood. the relative abundances of dinoflagellate cysts are very high, forming up to 77% of the total palynomorph assemblage, compared to 0–3% in the back-barrier sediments in gert-1 and gert-2 (johannessen et al. 1996). the coarsening-upwards to fining-upwards sandstones are intensely bioturbated, often to such a degree that primary sedimentary structures are obliterated. only 382 fig.10 4293 m core depth cl si sand pbl o ph io m or ph a– sk ol ith os ic hn of ab ri c gwen-2 h st h el m in th op sis –c ho nd rit es ic hn of ab ri c lo w er s ho re fa ce h el m in th op sis –t ei ch ic hn us –t ha la ss in ol de s– o ph io m or ph a ic hn of ab ri c m id dl e sh or ef ac e mtse2 sb2 h el m in th op sis –t ei ch ic hn us – th al as sin oi de s– o ph io m or ph a ic hn of ab ri c l. sh f. m.s. t st gamma ray 75 1500 0 50 40 30 20 10 m 60 gamma-ray 0 75 150 fig. 9. sedimentological core log of lower to middle shoreface sandstones from sequence 1 and the lower part of sequence 2 in the gwen-2 well on the gertrud plateau. the sandstones display several coarsening-upwards parasequences. for legend, see fig. 6. l.shf., lower shoreface; m.s., middle shoreface. 383 top base hel. gwen-2 core 2 box 10 core 2 box 11 core 2 box 12 core 2 box 9 core 2 box 13 tei. top parasequence oph. pal. oph. pla. sko. oph. tha. ? oph. sko. tha./hel. 10 cm fig. 10. middle shoreface parasequence in sequence 1 in the gwen-2 well. note the vertical succession of trace fossils within the parasequence: helminthopsis isp. (hel.) and thalassinoides isp. (tha.) in the lower levels and ophiomorpha isp. (oph.) in the upper part. for position of core photos, see fig. 9; for legend, see fig. 7. additional trace fossils: pal., palaeophycus heberti; pla., planolites isp.; sko., skolithos isp.; tei., teichichnus isp.; tha./hel., thalassinoides isp. burrow reworked by helminthopsis isp. 384 ravn-1 m 60 50 40 30 20 10 0 pblsicl 4155 m core depth fig. 12 h el m in th op sis –t ha la ss in oi de s ic hn of ab ri c t st h st o ph io m or ph a– th al as sin oi de s ic hn of ab ri c h el m in th op sis –a st er os om a ic hn of ab ri c t st sb2 h st mfs1 lo w er s ho re fa ce lo w er s ho re fa ce m id dl e sh or ef ac e lo w er s ho re fa ce mtse2 sb2 mtse2 m id dl e sh or ef ac e sand gamma-ray 0 75 150 fig. 11. sedimentological core log of lower to middle shoreface sandstones from the upper part of sequence 0, sequence 1 and the lower part of sequence 2 in the ravn-1 well, located on the heno plateau. the sandstones are represented by several coarseningupwards parasequences. for legend, see fig. 6. a few poorly-preserved cross-laminations are seen (figs 9, 11). a vertical suite of ichnofabrics characterises the sandstones. the helminthopsis–chondrites ichnofabric is typical of the lowermost sandy siltstones of the cored section in gwen-2 (fig. 9). the helminthopsis–teichichnus–thalassinoides–ophiomorpha ichnofabric occurs higher up in the succession in very fine-grained sandstones. the ophiomorpha–skolithos ichnofabric characterises the two parallel-laminated, medium-grained sandstones, 3 and 2 m thick, in the uppermost part of the coarsening-upwards succession, immediately below the conglomerate of gwen-2. the sandstone succession above the conglomerate is characterised by the helminthopsis–teichichnus–thalassinoides–ophiomorpha ichnofabric. the overall coarsening-upwards and fining-upwards sandstone succession of the heno formation may itself be constructed of several smaller-scale coarseningupwards (‘cleaning-upwards’) units, 1.5–10 m thick (figs 9, 10). such small-scale rhythmic units may also show a characteristic vertical succession of ichnofabrics. a coarsening-upwards unit (4.75 m thick) of very fineto fine-grained sandstones in the gwen-2 well is illustrated here as an example (fig. 10). helminthopsis isp. dominates at the base, followed upwards by thalassinoides isp. which is partly reworked by helminthopsis isp. the middle part of the unit is characterised by skolithos isp. and thalassinoides isp. (fig. 10). the upper, most coarse-grained part of the unit is dominated by many ophiomorpha isp. burrows and a few palaeophycus isp. traces. immediately above the top of the unit, clay-rich very fine-grained sandstones are characterised by planolites isp. and teichichnus isp. and form the base of the next coarsening-upwards, ‘cleaningupwards’ unit. out-sized quartz clasts, 0.5–3 mm in diameter, wellrounded and spherical, occur scattered within the bioturbated sandstones (fig. 9). the conglomerates are mostly matrix-supported, but a few 2–4 cm thick clastsupported conglomerate beds occur in jeppe-1 and gwen-2. the clasts are 0.5–2 cm in diameter, and the matrix consists of fineto medium-grained sandstone with pyrite, coal fragments and bivalve shells. interpretation the high diversity and high density trace fossil assemblage and the high relative abundance of marine palynomorphs, together with the land-derived organic matter, indicate that the sandstones and conglomerates were deposited on a marine shoreface with normal salinities and a strong terrestrial influence. the helminthopsis–chondrites ichnofabric is the result of deposit-feeding animals and indicates poor oxygen conditions (bromley 1990). the diverse trace fossil assemblage of the helminthopsis–teichichnus– thalassinoides–ophiomorpha ichnofabric was created by organisms that required more oxic conditions, indicating that the oxygen level at the sea bottom had increased (bromley 1990). thalassinoides and ophiomorpha traces represent permanent dwelling burrows, indicating that the energy conditions were higher such that the organisms constructed protective domiciles. the ichnofabrics of the lower and middle part of the coarsening-upwards succession represent the cruziana ichnofacies (seilacher 1967). the two parallel-laminated sandstone beds immediately below the conglomerate in gwen-2 (fig. 9) are characterised by a few ophiomorpha isp. and skolithos isp. burrows. this ichnofabric consists only of permanent dwelling burrows and high depositional energy is also indicated by the parallel lamination. the burrows represent dwelling structures and/or the activity of suspension-feeding animals and are typical of the skolithos ichnofacies (seilacher 1967). sedimentation rates were either sufficiently high or the degree of bioturbation was reduced, allowing preservation of the lamination. the sandstones were probably deposited in upper flow regime conditions during storms. the cruziana ichnofacies characterises sediments deposited between fair-weather wave base and storm wave base, within the lower shoreface (frey & pemberton 1985). the skolithos ichnofacies is characteristic of well-sorted sands deposited under high-energy conditions with frequent erosion and reworking, conditions typical of the middle to upper shoreface (pemberton et al. 1992). the few preserved cross-laminations indicate that traction currents operated over the sea floor giving rise to migrating small-scale ripples. the bioturbated sandstones characterising the bulk of the heno formation are interpreted as having been deposited by storm-generated currents that transported sand from the beach to the middle and lower shoreface. the scattered outsized matrix-supported quartz clasts were possibly deposited on scour surfaces by storm currents that swept across the sea floor; the clasts were subsequently dispersed in the sediment by burrowing. open burrows filled with shell debris and large clasts transported by storm-generated currents, as reported from comparable facies in the fulmar formation (taylor & 385 gawthorpe 1993), have not been recognised in the shoreface sandstones of the heno formation. the fact that the shoreface sandstones of the heno formation are completely bioturbated suggests that each storm sand bed was relatively thin and/or that the mid-tier and shallow-tier trace fossils were able to completely penetrate the storm sand beds. the coarsening-upwards succession represents progradation of a lower to middle shoreface sandstone wedge, while the succeeding fining-upwards succession represents the subsequent backstepping of middle to lower shoreface sandstone wedges. the small-scale coarsening-upwards units, 1.5–10 m thick in gwen-2 (fig. 9), possibly represent minor shoreface progradational events. the coarsening-upwards unit illustrated from the gwen-2 well (fig. 10), shows ichnofabrics of the cruziana ichnofacies at the base, suggesting moderate energy conditions, and the skolithos ichnofacies at the top, suggesting high-energy conditions (seilacher 1967). the very fine-grained, clayey sandstones that abruptly overlie the coarsening-upwards unit and are dominated by planolites and teichichnus traces produced by deposit feeders, indicate a dramatic change to much lower energy conditions (fig. 10). in general, the vertical successions of ichnofabrics in the heno formation are closely comparable to those described by taylor & gawthorpe (1993) from the upper kimmeridgian – lower volgian shoreface sandstones of the ula formation of the southern norwegian central graben. in contrast, however, neither the siphonichnus ichnofabric nor the glossifungites ichnofacies, which have been described from the ula and fulmar formations (taylor & gawthorpe 1993), were recognised in the heno formation. the conglomerates that abruptly overlie the coarsening-upwards shoreface sandstones and are overlain by fining-upwards middle shoreface sandstones are interpreted as lags formed during transgression upon a marine transgressive surface of erosion, mtse2 (fig. 9). the clasts of the conglomerates are much larger than the out-sized clasts in the shoreface sandstones. the conglomerates are therefore interpreted to represent storm events on the beach/shoreface or reworked fluvial deposits formed during maximum regression. shoreface erosion during the subsequent transgression is thought to have obliterated evidence of individual coarse-grained storm beds and evidence of subaerial exposure. the large number of coal clasts within the conglomerates, however, may indicate fluvial erosion of peat deposits. the backstepping upper part of the shoreface sandstone succession is abruptly overlain by offshore claystones of the farsund formation, a feature also seen in the fulmar formation and, to some degree, in the ula formation (howell et al. 1996). further occurrences regressive to transgressive shoreface sandstones and conglomerates occur in all the wells on the heno and gertrud plateaus with the exception of diamant-1 where only transgressive sandstones occur and gert-2 where conglomerates are missing (fig. 8). the section in the ravn-1 well, in the southern part of the heno plateau (figs 1, 4, 5, 8), includes two pebble conglomerate beds in the more than 100 m thick fine-grained sandstone succession of the heno formation (fig. 11). the lower conglomerate is 0.25 m thick, with clast sizes of up to 3 cm in diameter; no imbrication of clasts is seen. the upper conglomerate is c. 2 m thick and varies from clastto matrix-supported, with clast sizes of 0.3–1 cm (fig. 12); it overlies fineto mediumgrained sandstones characterised by an ophiomorpha– thalassinoides ichnofabric (fig. 11). an overall finingupwards succession characterises the section above both of the conglomerate units in ravn-1 (figs 8, 11). regressive–transgressive shoreface sandstones without conglomerates are present in all wells in the feda graben and the salt dome province adjacent to the plateau areas and on the deeper, south-eastern part of the heno plateau (figs 5, 8, 13). the grain size becomes finer as the distance to the plateau areas increases, suggesting that shoreface sandstones are only represented close to the margins of the grabens. the sandstones show a gradual fining-upwards and grade up into offshore claystones of the farsund formation. the lower shoreface sandstones in the salt dome province, on the south-east heno plateau and the southern part of the gertrud plateau are described below. shoreface sandstones are not observed in the tail end graben (fig. 13). sharp-based lower shoreface sandstones description in the elly-2 well, in the deeper, south-eastern part of the heno plateau, very fine-grained clayey sandstones form a succession 185 m thick characterised by relatively low gamma-ray values (figs 1, 4, 8). the sandstones abruptly overlie claystones with high gamma-ray values, and are overlain by similar claystones. the sand386 387 o ph . o ph . o ph . o sq c m t se 2 sb 2 r av n1 t o p b as e c or e 3 bo x 17 c or e 3 bo x 18 c or e 3 bo x 19 c or e 3 bo x 20 te r. c or e 3 bo x 21 c or e 3 bo x 22 c or e 3 bo x 23 b c or e 4 bo x 1 c or e 4 bo x 2 c or e 4 bo x 3 as t. o ph . th a. ve ry fi ne -g ra in ed sa nd st on e 10 c m sh ar pba se d fin egr ai ne d sa nd st on e fi g. 1 2. l o w er t o m id d le s h o re fa ce s an d st o n es a n d c o n gl o m er at es i n t h e r av n -1 w el l; th e sb 2 se q u en ce b o u n d ar y is p la ce d a t th e b as e o f th e co n gl o m er at e. n o te t h e sh ar p -b as ed , fi n egr ai n ed s an d st o n es c o n ta in in g o u tsi ze d q u ar tz c la st s (o sq c ) o ve rl yi n g ve ry f in egr ai n ed s an d st o n es . fo r p o si tio n o f th e co re p h o to , se e fi g. 1 1; f o r le ge n d , se e fi g. 7 . t ra ce fo ss ils : a st ., a st er os om a is p .; o p h ., o ph io m or ph a n od u sa ; t er ., t er eb el li n a is p .; t h a ., t h a la ss in oi d es is p . su rf ac es : m t se , m ar in e tr an sg re ss iv e su rf ac e o f er o si o n ; sb , se q u en ce b o u n d ar y. stones and the claystones contain abundant dinoflagellates. two units (0.5–1 m thick) in the claystones immediately below the sandstones show low gammaray and sonic readings and high resistivity readings and form discrete log markers; cuttings samples suggest that these units consist of carbonate-cemented sediment. these log markers are also recognised in the elly-2, falk-1, skarv-1 and edna-1 wells on the south-eastern heno plateau and in the salt dome province (fig. 8). the sandstones sharply overlie the uppermost calcareous marker bed in elly-2 whereas in the other wells, the uppermost calcareous bed is situated 10–25 m below the sharp-based, very fine-grained clayey sandstones. a 50 m thick unit in the middle part of the sandstone succession in elly-2 shows slightly higher gamma-ray readings, indicating that it is more clay-rich than the sandstones above and below (fig. 8). a core (20 m thick) in the lowermost part of this section (fig. 8) comprises very fine-grained, glauconitic sandstones and sandy siltstones, which are heavily bioturbated with a high diversity trace fossil assemblage. in the lower to middle part of the core, mostly comprising sandy siltstones, the ichnofauna consists only of helminthopsis isp. in the more sandy intervals, the diversity of trace fossils is much higher with burrows of asterosoma isp., chondrites isp., helminthopsis horizontalis, ophiomorpha isp., palaeophycus isp., planolites isp., rhizocorallium isp., skolithos isp., teichichnus isp., terebellina isp. and thalassinoides isp., defining the cruziana ichnofacies (seilacher 1967). in the upper half of the core, several very fine-grained sandstone beds, 8–10 cm thick, are interbedded with clayey to silty, very fine-grained sandstones. no primary sedimentary structures are preserved in the sandstones, presumably due to the intensive bioturbation. bed boundaries are also thoroughly bioturbated. interpretation the high diversity trace fossil assemblage and the large number of dinoflagellate cysts indicate that the finegrained sandstones were deposited under well-oxygenated conditions in a fully marine environment, probably the lower shoreface. the sandy siltstones with an ichnofabric consisting solely of helminthopsis isp. were probably deposited under poorly oxygenated conditions (bromley 1990; johannessen 1995). the very fine-grained sandstones, interbedded with clayey to silty, very fine-grained sandstones in the upper part of the elly-2 section, were probably deposited from stormgenerated currents. formation of glauconite takes place below wave base (odin & fullagar 1988; hansley & whitney 1990). from their studies of the upper cretaceous shannon sandstone (usa), walker & bergman (1993) suggested that glauconite sands originally formed below wave base but were subsequently reworked during a forced regression into wave-dominated shoreface sandstones. the abnormally high gamma-ray values in the lower shoreface sandstones are due to the large content of glauconite which contains radioactive potassium. the sharp-based sandstones overlying the offshore claystones in elly-2 indicate a sudden increase in depositional energy. the uppermost calcareous marker bed is situated immediately below the sharp-based sandstone, in contrast to the other wells in the south-east of the heno plateau and the salt dome province where the calcareous marker occurs some 10–25 m below the sharp-based sandstones. it is considered likely therefore that up to 25 m of offshore claystones were stripped away prior to deposition of the sandstones in the elly-2 area. comparable shoreface sandstones resting on a sharp, erosional surface have been described by plint (1988) from the cardium formation in canada. further occurrences sharp-based, clayey, very fine-grained lower shoreface sandstones are also recognised in the falk-1, skarv-1 and edna-1 wells situated on the south-east heno plateau and in the salt dome province, and also in the gwen-2 and q-1 wells, on the southernmost part of the gertrud plateau (fig. 8). the sandstones of these wells are finergrained than the sandstones in elly-2, based on cuttings samples and petrophysical logs. a casing shoe was placed at the bottom of the sharp-based sandstones in the elly-2 and edna-1 wells, such that the abrupt nature of the boundary may be in part an artefact. similar sandstones occur, however, in the falk-1 and skarv-1 wells where no casing shoe disturbs this boundary (fig. 8); the sharp-based sandstones observed in these wells are also interpreted to have been deposited on an erosional surface. sharp-based shoreface sandstones have not been recorded from the tail end graben (fig. 13). depositional sequences and basin development high resolution sequence stratigraphic interpretation of deposits in an active rift basin such as the central 388 graben can be problematic for a number of reasons. subsidence rates in different parts of the basin may vary strongly. local sand sources may be shut off almost instantaneously due to drowning as a result of relative sea-level rise (johannessen et al. 1996). footwall uplift may create new sediment source areas whilst subsidence of hanging-walls can create increased accommodation space resulting in thick prograding shoreface wedges (gawthorpe et al. 1994; johannessen et al. 1996). furthermore, small rotating half-grabens near the main sediment sources may trap sediment before reaching the deeper part of the basin. the kimmeridgian sandstone succession of the danish central graben is subdivided into two sequences (sequences 1 and 2) that are described below. the underlying sequence (sequence 0) is largely outside the scope of this study but is also described briefly below. sequences 0, 1 and 2 of this study are equivalent to the kimm-1, kimm-2 and kimm-3 sequences, respectively, in the regional jurassic sequence stratigraphic study by andsbjerg & dybkjær (2003, this volume). sequence 0 as noted above, sequence 0, beneath the heno formation sandstones, is outside the scope of the study but the upper levels are discussed here to provide the stratigraphic context of the succeeding sand-rich sequences; for further details of this sequence, the reader is referred to andsbjerg & dybkjær (2003, this volume) where sequence 0 of this study corresponds to their kimm-1 sequence. the upper part of sequence 0 consists of a highstand systems tract overlying a maximum flooding surface (mfs0); it is present on the south-eastern part of the gertrud plateau, the south-eastern part of the heno plateau and in the tail end graben (figs 5, 8, 13). highstand systems tract the coarsening-upwards succession above the lola formation offshore claystones, which span the mfs0 (figs 14a, 15), is interpreted as a progradational shoreface wedge and is referred to the highstand systems tract (figs 5, 8, 13). the most coarse-grained shoreface sandstones occur in the w-1, ravn-1 and ravn-2 wells on the southern part of the heno plateau (figs 5, 8). in ravn-1, three well-defined parasequences form a forestepping parasequence set. in the gwen-2 and q-1 wells, on the southern part of the gertrud plateau, the highstand systems tract consists of weakly coarsening-upwards offshore claystones (fig. 5). a succession of offshore claystones grading up 389 4266 m b. msl 3246 m b. msl sb0 hst hst tst tst 100 m 50 km danish central graben nora-1 ravn-1 tail end grabenheno plateau sb1 sb2+ mtse2 mfs0 mfs1 ravn-1 gr nora-1 gr fig. 13. gamma-ray log correlation from the easternmost part of the heno plateau (ravn-1) out into the tail end graben (nora-1), adjacent to the heno plateau. sharp-based shoreface sandstones are not observed in the tail end graben. sequence stratigraphic boundaries in the nora-1 well are from andsbjerg & dybkjær (2003, this volume). for legend, see fig. 5. 390 ? ? ?? a c db ea rl y k im m er id gi an ba yl ei c hr on oz on e m fs 0 lo la f or m at io n ea rl y k im m er id gi an c ym od oc e– m ut ab ili s c hr on oz on e tr an sg re ss io n lo la f or m at io n an d h en o fo rm at io n 20 k m 20 k m 20 k m 20 k m ea rl y k im m er id gi an c ym od oc e c hr on oz on e sb 1 lo la f or m at io n an d h en o fo rm at io n la te k im m er id gi an m ut ab ili s c hr on oz on e m fs 1 lo la f or m at io n an d h en o fo rm at io n 391 e g f o ffs ho re c la ys to ne s lo w er s ho re fa ce c la ye y sa nd st on e m id dl e to u pp er s ho re fa ce s an ds to ne sh or ef ac e co ng lo m er at e ba ck -b ar ri er s ed im en ts er os io n/ no nde po si tio n c la st ic s up pl y r eg io na l d ra in ag e la te k im m er id gi an m ut ab ili s c hr on oz on e ea rl y tr an sg re ss io n lo la f or m at io n an d h en o fo rm at io n 20 k m 20 k m 20 k m la te k im m er id gi an m ut ab ili s c hr on oz on e sb 2 lo la f or m at io n an d h en o fo rm at io n la te k im m er id gi an m ut ab ili s c hr on oz on e m aj or fl oo di ng of h en o fo rm at io n fi g. 1 4. s er ie s o f p al ae o ge o gr ap h ic m ap s ill u st ra tin g b as in e vo lu tio n d u ri n g th e k im m er id gi an . t h e m id d le t o u p p er s h o re fa ce s an d st o n es i n d ic at ed o n a ll th e m ap s al o n g th e r in gk ø b in g– fy n h ig h f ri n gi n g th e sø gn e b as in , th e t ai l e n d g ra b en a n d t h e sa lt d o m e p ro vi n ce a re n o t d o cu m en te d b y w el ls b u t ar e in fe rr ed o n t h e b as is o f th e re gi o n al p al ae o ge o gr ap h ic r ec o n st ru ct io n . d at a fr o m t h e n o rw eg ia n a n d u k s ec to rs a re fr o m b er ga n e t a l. (1 98 9) , r at te y & h ay w ar d ( 19 93 ), r ic h ar d s et a l. (1 99 3) a n d m ac ke rt ic h ( 19 96 ). into siltstones occurs on the south-east heno plateau and in the salt dome province (falk-1, elly-2; figs 8, 13). in contrast, in the skarv-1 well, a coarsening-upwards trend is not evident; the offshore claystones and siltstones are abruptly overlain by clayey, very fine-grained sandstones in this well (fig. 8). sequence 0 is not recognised in the feda graben nor on the northern part of the gertrud plateau indicating either uplift and erosion or non-deposition due to the positive nature of these structures (figs 14a, b, 15, 16a, b). the shoreface sandstones prograded towards the east across the heno plateau and probably the easternmost gertrud plateau before pinching out in the westernmost tail end graben and søgne basin (fig. 14b). shoreface sandstones may also have been deposited along the easternmost flank of these basins, bordering the ringkøbing–fyn high. sequence 1 in the south-eastern part of the gertrud plateau and the tail end graben, sequence 1 consists of lowstand, transgressive and highstand systems tracts (figs 8, 13). on the northern part of the gertrud plateau, the southern part of the heno plateau and in the feda graben, sequence 1 consists of transgressive and highstand systems tracts (figs 8, 13). on the western part of the heno plateau, sequence 1 consists only of sediments of the transgressive systems tract (fig. 8). sequence 1 392 lo la f or m at io n fa rs un d fo rm at io n fo rm at io ns ba yl ei c ym od oc e m ut ab ili s eu do xu s a ut is si od or en si s age key surfaces key surfaces norwegian sector danish sector la te ju ra ss ic la te k im m er id gi an mfs1 sb2 mfs2 mfs0 mfs2 mfs1 sb2 sb1 mfs0 farsund formation sb1 lola formationea rl y k im m er id gi an ? ? ? salt dome provinceheno plateaugertrud plateau/grabenfeda graben nw se sb1 sb0 sb1 sb0 sb0 sb0 ? shoreface conglomerate middle–upper shoreface sandstone hiatus back-barrier sediments lower shoreface clayey sandstone offshore claystone erosion surface 2/11-7 gert-4 gert-1 2/12-1 gert-2 jeppe-1 gwen-2 q-1 w-1 ravn-1 elly-1 edna-1 heno formation fig. 15. late jurassic time stratigraphic correlation chart of the northern part of the danish central graben and the southernmost part of the norwegian central graben. the question-marks in the norwegian part of the section indicate that lower shoreface sandstones and back-barrier sediments of early kimmeridgian age have not been documented to date, but are predicted on the basis of this study. corresponds to the kimm-2 sequence of andsbjerg & dybkjær (2003, this volume). lower sequence boundary (sb1) and the lowstand systems tract the sequence boundary sb1 merges with the regional ‘base upper jurassic’ unconformity in the western part of the heno plateau, the northern part of the gertrud plateau and the feda graben (figs 5, 8). this suggests that these areas experienced sediment bypass and that sediment was deposited farther out in the deeper parts of the basin (figs 14b, 15, 16b, c). the conglomerates in ravn-1 and ravn-2 were probably deposited during sea-level fall and subsequent sea-level rise and the base of the conglomerate unit is thus interpreted as a sequence boundary, sb1 (figs 5, 8). in the w-1 well, the sb1 is placed at the top of the regressive sandstones, i.e. at the maximum regression. the sharp-based sandstones of the southern part of the gertrud plateau (q-1, gwen-2) and the south-east heno plateau and salt dome province (falk-1, elly-2, skarv1, edna-1) are situated on a regional erosional surface caused by a fall in sea level. the base of the sharp-based sandstones thus represents a sequence boundary (sb1; figs 5, 8). due to this inferred fall in sea level, the sharp-based sandstones extend farther out in the basin than the shoreface sandstones of the underlying highstand systems tract (fig. 15). it is thought likely that comparable sharp-based sandstones were deposited in the feda graben, but well data are presently not available. as the sharp-based sandstones shale out over short distances basinward, they probably represent the most distal part of a progradational wedge of shoreface sandstone. this suggests that they were deposited close to the time when relative sea level reached its lowest position; they are thus referred to the lowstand systems tract (posamentier & vail 1988; van wagoner et al. 1990; hunt & tucker 1992, 1995; helland-hansen & gjelberg 1994). on the south-east heno plateau and in the salt dome province, the lowstand systems tract consists of a package (5–10 m thick) of aggradational clayey, very finegrained, lower shoreface sandstones. on the south-eastern part of the gertrud plateau, the lowstand systems tract consists of lower shoreface siltstones, forming a unit 5–10 m thick (figs 8, 13). the top of the lowstand systems tract is defined at the base of a fining-upwards succession, which is referred to the succeeding transgressive systems tract (figs 14b, 15, 16b). sharp-based shoreface sandstones are not seen in the tail end graben (fig. 13). transgressive systems tract and mfs1 the transgressive systems tract is recognised in all the wells in the study area. the transition from the lowstand systems tract to the transgressive systems tract is similarly developed on the southern part of the gertrud plateau, on the south-east heno plateau and in the salt dome province (figs 8, 13). in the gwen-2 well, a weakly fining-upwards succession of clayey siltstones, c. 40 m thick, deposited on a lower shoreface, shows a backstepping pattern capped by a maximum flooding surface (mfs1; fig. 5). in the ravn-1 well on the heno plateau, the basal sequence boundary (sb1) is overlain by a conglomerate, 0.25 m thick, followed by three backstepping parasequences, each 2–5 m thick, representing the transgressive systems tract (figs 8, 13). in the feda graben (2/12-1) and on the gertrud plateau (gert-2; fig. 5), the transgressive systems tract consists of a vertical succession of back-barrier sediments, shoreface sandstones and offshore claystones; the tract is c. 180 m and c. 25 m thick, respectively, in the 2/12-1 and gert-2 wells (johannessen et al. 1996). the transgressive systems tract in the feda graben and the northern part of the gertrud plateau overlies sequence boundary sb1, which here coincides with the ‘base upper jurassic’ unconformity (fig. 5). in the diamant-1 well, near the inge high on the westernmost part of the heno plateau, a succession of backbarrier sediments, 23 m thick, makes up part of the transgressive systems tract (fig. 8). this succession overlies the sequence boundary sb1, coinciding with the ‘base upper jurassic’ unconformity (fig. 8). the back-barrier sediments are erosionally overlain by a conglomerate, 2 m thick. the conglomerate in the diamant-1 well was probably deposited during a fall in relative sea level, related to the formation of sb2, and was later reworked during a rise in relative sea level associated with the formation of a marine transgressive surface of erosion (mtse2). the base of the conglomerate thus represents the mtse1, sb2 and mtse2 merged together, defining the top of the transgressive systems tract (fig. 8). in the eg-1 well, the transgressive systems tract consists of a fining-upwards succession (10 m thick) of very fine-grained lower shoreface sandstones (fig. 8). a thin conglomerate, 8 cm thick, abruptly overlying the shoreface sandstones, defines the upper sequence 393 394 c fs sb 1 m fs 1 m fs 0 sb 0 sb 1 m fs 0 sb 0 m fs 0 sb 0 a b w -1 r av n1 d ia m an t1 ed na -1 el ly -2 g w en -2 q -1 je pp e1 g er t4 g er t1 g er t2 2/ 12 -1 2/ 11 -7 m fs 1 sb 0 + sb 1 sb 0 + sb 1 eg -1 fe da g ra be n g er tr ud p la te au h en o pl at ea u sa lt d om e pr ov in ce 395 e sb 1 d sb 2 m fs 1 fsfsfs m fs 0 sb 0 sb 1 sb 2 m fs 1 m fs 0 sb 0 tr an sg re ss iv e sy st em s tr ac t h ig hs ta nd s ys te m s tr ac t lo w st an d sy st em s tr ac t sb 2 m fs 1 sb 0 + sb 1 fssb 2 m fs 1 sb 0 + sb 1 fs fl oo di ng s ur fa ce m fs 1 m ax im um fl oo di ng s ur fa ce , s eq ue nc e 1 sb 1 se qu en ce b ou nd ar y, se qu en ce 1 fi g. 1 6. s eq u en tia l b as in e vo lu tio n d u ri n g th e k im m er id gi an , d ep ic te d b y ga m m ara y lo g p an el s fr o m t h e fe d a g ra b en , g er tr u d p la te au , h en o p la te au a n d s al t d o m e p ro vi n ce ; th e se q u en ce s tr at ig ra p h ic s u rf ac es s b 0, m fs 0, s b 1, m fs 1, s b 2 an d f s (s eq u en ce 2 ) ar e u se d a s d at u m l in es . n o te t h at t h e gr ab en a re as a re d o m in at ed b y th ic k h ig h st an d a n d t ra n sg re ss iv e sy st em s tr ac ts c o m p ar ed t o t h e p la te au a re as . t h e p la te au a re as n ea re st t o t h e se d im en t so u rc e ar ea , as e xe m p lif ie d b y th e d ia m an t1 an d e g1 w el ls n ea r th e in ge h ig h , ar e d o m in at ed b y th in t ra n sg re ss iv e sy st em s tr ac ts . t h e sh o re fa ce s an d st o n es d ep o si te d i n t h e sh al lo w p la te au a re as a re g en er al ly a b ru p tly o ve rl ai n b y o ff sh o re c la ys to n es . t h e sh o re fa ce s an d st o n es d ep o si te d i n t h e d ee p p la te au a n d g ra b en a re as a re m o re f in egr ai n ed a n d g ra d u al ly f in e u p w ar d s to o ff sh o re c la ys to n es . n o te t h at t h e sh o re fa ce s an d st o n es o n t h e g er tr u d p la te au b ec o m e m o re f in egr ai n ed i n t h e fe d a g ra b en a n d s h al e o u t in 2 /1 17. boundary (sb2); middle shoreface sandstones were probably eroded during the fall in sea level that created this sequence boundary. a marine transgressive surface of erosion, mtse1, is inferred to separate transgressive back-barrier sediments in diamant-1 from the transgressive shoreface sandstones observed in eg-1 (fig. 8). the heno and gertrud plateaus were transgressed both from the east and the west (figs 14b–d, 15, 16b, c). highstand systems tract and the upper sequence boundary (sb2) the highstand systems tract consists of shoreface sandstones in all the wells in the study area, with the exception of diamant-1 and eg-1 in the westernmost part of the heno plateau, where deep erosion at sb2 is indicated. the highstand systems tract on the gertrud plateau is characterised by a coarsening-upwards succession up to 70 m thick of intensely bioturbated siltstones and fineto medium-grained sandstones, deposited on the lower to middle shoreface (gert-2, jeppe-1, gwen-2, q-1; fig. 5). the top of the systems tract is defined at the most coarse-grained part of the sandstones in gert-2, and at the base of the conglomerate that caps the sandstones in the jeppe-1, gwen-2 and q-1 wells (figs 5, 9, 10). the conglomerate was probably deposited during relative sea-level fall and subsequent sea-level rise, as described earlier, and the base of the conglomerate is thus interpreted as a sequence boundary (sb2; fig. 5). similarly, the highstand systems tract in the ravn-1 well, on the heno plateau, is composed of a fine-grained, bioturbated shoreface sandstone unit, 2.3 m thick, capped by a 2 m thick conglomerate, the base of which defines sb2 (figs 8, 11, 12). in the w-1 well, the top of the highstand systems tract is placed at the base of a conglomerate, the presence of which is inferred from the low gamma-ray readings and cuttings samples (fig. 8). in the basin areas (tail end graben, south-eastern heno plateau, salt dome province), the highstand systems tract is 50–200 m thick and consists of a coarsening-upwards succession of clayey sediments. this suggests that the graben areas subsided at a higher rate than when the lowstand sediments of sequence 1 were deposited. locating the upper sequence boundary (sb2) in the basin areas is problematic. the rate of graben subsidence may have exceeded the rate of eustatic fall, such that sequence boundary sb2 is a correlative conformity; the sequence boundary would thus be placed at the most coarse-grained part of the lower shoreface sandstones, marking the maximum regression (figs 8, 13, 17a). alternatively, if the shoreface sandstones in the basin areas were deposited during a relative fall in sea level, as suggested on the plateau areas, part of the sandy succession in edna-1 may represent sediments of the lowstand systems tract (fig. 17b). following this line of reasoning, the sequence boundary should be placed beneath, or in the lowermost levels, of the coarsening-upwards shoreface sandstones (fig. 17b). unfortunately, the biostratigraphic, seismic and petrophysical data are ambiguous with respect to these two alternative interpretations; in this study, the first interpretation is preferred, i.e. the thick coarsening-upwards succession of shoreface sandstones is referred to the highstand systems tract (figs 8, 17a). the highstand systems tract in the feda graben is also characterised by a thick coarsening-upwards succession, 75–100 m thick, without notable breaks (fig. 5). the lowermost part of the systems tract is more finegrained in the feda graben than in the salt dome province; it comprises a basal unit of organic-rich offshore claystones overlain by claystones and siltstones, followed by very fine-grained lower shoreface sandstones (fig. 5). in the norwegian well, 2/12-7, farther out in the feda graben, the correlative section is 110 m thick and is composed of regressive offshore claystones and siltstones. significant stratigraphic breaks have not been recognised and conglomerates are absent from the succession. the upper sequence boundary (sb2) is placed at the most coarse-grained part of the lower shoreface sandstones, marking the maximum regression (fig. 5). the conformable nature of sb2 in this area suggests that the subsidence rate of the feda graben was higher than the rate of eustatic sea-level fall such that fluvial/beach conglomerates did not prograde into the graben. the conglomerates of the gertrud plateau (jeppe-1 and gwen-2) are thin and relatively fine-grained and probably represent a smaller fall in sea level than that suggested by the ravn-1 succession, where low rates of subsidence (or even uplift) may have enhanced the effect of a small eustatic sea-level fall. conversely, high rates of subsidence on the gertrud plateau may have obscured significant eustatic sea-level fall. the shoreface sandstones prograded towards the east and north-east across the heno plateau and towards the west and south-west on the gertrud plateau (figs 14e, 15; 16d). 396 sequence 2 in sequence 2, sandstones are only found in the transgressive systems tract (figs 5, 8); the highstand systems tract consists solely of offshore claystones and shoreface siltstones and is not described further in this paper; for discussion of this succession, the reader is referred to johannessen et al. (1996) and andsbjerg & dybkjær (2003, this volume). sequence 2 corresponds to the kimm-3 sequence of andsbjerg & dybkjær (2003, this volume). transgressive systems tract sediments of the transgressive systems tract are distributed throughout the study area. on the southern heno plateau, the sandstones of the transgressive systems tract form an aggrading to backstepping parasequence set (55 m thick in ravn-1) composed of middle and lower shoreface sandstones; these are abruptly overlain by offshore claystones of the farsund formation (figs 8, 11, 12). the base of the offshore claystones represents a major flooding surface which can be traced 397 lstlst fs tst tst hst hst hst sb0 sb2+ mtse2 tst tst lst hst sb0 fs sb2+ mtse2 sb1 mfs0 mfs1 sb1 sb2 mfs0 mfs1 fsfs fs plateau ravn-1 gr basin edna-1 gr a b 100 m basin edna-1 gr plateau ravn-1 gr fig. 17. two alternative interpretations concerning the correlation of sequence boundary sb2 from the heno plateau (ravn-1) to the salt dome province (edna-1); for legend, see fig. 5. a: a fall in sea level is indicated on the heno plateau by the 2 m thick conglomerate that erosionally overlies fine-grained shoreface sandstones in the ravn-1 well. no indication of this fall in sea level is evident in the basin (e.g. edna-1). consequently, the sequence boundary (sb2) is correlated from the base of the conglomerate on the heno plateau to the most coarse-grained part of the lower shoreface sandstones, marking maximum regression in the basin. b: alternatively, as discussed in the text, it can be argued that the correlative conformity (sb2) should be placed at some point in the lower part of the coarsening-upwards shoreface sandstones. according to this interpretation, the coarsening-upwards lowstand systems tract represents sediments that bypassed the plateau areas. all over the gertrud plateau and the northern, shallow part of the heno plateau (figs 5, 8). this suggests that the plateaus were abruptly flooded, either due to a rapid rise in relative sea level or as a result of an abrupt shut-down in coarse clastic supply, due to submergence of highs that acted as sand source areas (johannessen et al. 1996; howell et al. 1996). over much of the gertrud plateau, the transgressive systems tract consists of backstepping middle shoreface sandstones (10–28 m thick) overlying the basal conglomerate (e.g. gwen-2, q-1, jeppe-1; fig. 5). the sandstones are directly overlain by offshore claystones of the farsund formation in all three wells and lower shoreface sandstones are not recognised, in contrast to wells of the heno plateau (figs 5, 8). in the gert-2 well, however, on the westernmost part of the gertrud plateau close to the feda graben, an aggradational succession (43 m thick) composed of clayey, very fine-grained lower shoreface sandstones is referred to the transgressive systems tract (fig. 5). the sandstones are abruptly overlain by offshore claystones indicating an overall deepening. the greater thickness and finer grain size in gert-2 compared to the jeppe-1, gwen-2 and q-1 wells suggests that the westernmost part of the gertrud plateau experienced higher subsidence rates than the rest of the plateau and that gert-2 was situated more distally, relative to the other wells. the sedimentation rate in the westernmost part of the plateau must have been high to keep up with the continuously increasing accommodation space. this may have been the result of synsedimentary fault activity at the western border of the gertrud plateau (johannessen et al. 1996). in the feda graben, the transgressive systems tract consists of a fining-upwards succession of clayey very fine-grained, lower shoreface sandstones and offshore siltstones and claystones, as seen in the gert-4 (60 m thick), 2/12-1 (45 m) and 2/11-7 (15 m) wells (fig. 5). this fining-upwards transgressive succession is absent in the gert-1 well, in the easternmost feda graben (fig. 5); seismic data suggest that a normal fault may have cut out this part of the section in this well (johannessen et al. 1996). on the south-east heno plateau and in the salt dome province, clayey, very fine-grained, lower shoreface sandstones, 30–60 m thick, gradually fine upwards into offshore claystones (fig. 8). in the diamant-1 well, on the westernmost part of the heno plateau near the inge high, the uppermost portion (10 m) of the fineto medium-grained upper shoreface sandstone succession overlying the conglomerate shows an overall finingupwards, transgressive trend and is abruptly overlain by offshore claystones (fig. 8). these sandstones represent the most permeable (1.2–2.7 darcy) and most porous (22–30%) jurassic sandstones recorded in the danish central graben. in the tail end graben, the transgressive systems tract comprises fine-grained clay-rich sediments (50 m thick in nora-1); in the feda graben, this tract consists of a succession of clayey, very fine-grained sandstones and claystones, 80 m thick in gert-4 (figs 5, 8, 13). the transgressive systems tract is thick and coarser on the heno plateau in ravn-1 (55 m thick section of mediumto very fine-grained sandstones; fig. 8) but is thinner, (c. 25 m thick) over the remaining part of the heno plateau. on the gertrud plateau, the transgressive systems tract is thin (10 m thick, fine-grained sandstone). the thin nature of the transgressive systems tract on the plateau areas compared with the graben areas suggests that the platform experienced a lower subsidence rate. the heno and gertrud plateaus were transgressed from both the east and the west (figs 14e–g, 15, 16e). discussion the geometry of the depositional units and the interpreted palaeogeography indicate that the sandstones were sourced from the mandal high, north-east of the gertrud plateau, and the inge–mads highs west of the heno plateau, which formed part of the mid north sea high during the kimmeridgian. the gertrud and heno plateaus formed a contiguous area. two grabens fringed the platform areas: the feda graben towards the northwest and the tail end graben towards the south-east. graben versus plateau successions the identified sequences show major lateral differences from the plateaus to the grabens. on the shallow plateaus, the sequences are thin and more coarsegrained than in the adjacent grabens (figs 5, 8, 13). the conglomerates that mark sequence boundaries sb1 and sb2 on the shallow heno plateau and sb2 on the gertrud plateau are not observed in the graben areas and the deeper, south-eastern portion of the heno plateau, indicating that the sea level did not fall to such a degree that rivers or beaches reached these areas. this was probably due to the higher rates of subsidence in the two grabens relative to the plateau areas. 398 two major sea-level falls were responsible for the formation of the conglomerates that mark sequence boundaries sb1 and sb2 on the heno plateau. on the deep, south-east heno plateau and in the salt dome province, only sb1 is distinctly developed, being defined by the base of sharp-based lowstand shoreface sandstones. the succeeding sequence boundary sb2 is developed as a correlative conformity in the grabens. this suggests that the graben was subsiding at smaller rates during the fall in relative sea level responsible for sb1 than during the subsequent fall in relative sea level (sb2). the shoreface sandstones of the transgressive systems tract of sequence 2 in the graben areas gradually fine upwards and grade into the offshore claystones of the farsund formation. in contrast, shoreface sandstones on the plateau areas are abruptly overlain by offshore claystones indicating an abrupt cessation of supply of coarse sediment due to submergence of the source areas. the abrupt cessation of coarse sediment supply probably happened after deposition of the sandy sediments in the graben areas, since no abrupt decrease in grain-size is seen here. potential distribution of reservoir sandstones the sequence stratigraphic interpretation presented here has implications for the potential distribution of sandstone reservoirs elsewhere in the northern danish central graben, in areas that are presently undrilled. incised valley fill during sea-level fall, rivers probably intersected previously deposited shoreface sandstones resulting in sediment transport across the heno and gertrud plateaus and deposition basinward in the tail end and feda grabens, on the south-east heno plateau and in the salt dome province. during the subsequent transgression, high-energy shoreface processes removed all fluvial sediments on the plateaus, leaving only a transgressive lag behind, forming the most coarse-grained sediment resting upon the marine transgressive surfaces of erosion (mtse). where fluvial incision was sufficiently deep, however, incised valley fills may have been preserved, at least in part. such inferred incised valleyfills, possibly comprising thick fluvial and estuarine coarse-grained sandstones with good reservoir properties, may represent an important exploration target. lowstand shoreface sandstones the lowstand shoreface clayey sandstones of sequence 1 are only about 10 m thick on the south-east heno plateau and in the salt dome province. thicker, coarser and better-sorted shoreface sandstones may be found near the outer reaches of incised valleys, but have not yet been drilled. such sandstones may constitute good reservoirs and are situated closer to the source rocks of the farsund formation in the tail end graben than the sandstones upon the plateaus. the sea-level fall in the eastern part of the feda graben and the northern part of the gertrud plateau is represented by a bypass surface (sb1) amalgamated with the ‘base jurassic’ unconformity (fig. 15). lowstand shoreface sandstones are thus expected to be present farther out in the feda graben, as observed in the tail end graben. stranded parasequences between the plateau areas and the tail end graben and the south-eastern part of the heno plateau, stranded parasequences (the forced regressive systems tract of hunt & tucker 1992, 1995) comprising shoreface sandstones may have been deposited if the sea-level fall occurred in steps. such stranded parasequences may represent important future reservoir targets. outer rough basin the shoreface ‘fife sandstones’ of the fife field situated immediately north-west of the danish central graben in the uk sector are very similar to those of the heno formation. the ‘fife sandstones’ have been referred to the early–middle volgian (mackertich 1996) although a recent study suggested that they may extend down to the latest kimmeridgian (spathopoulos et al. 2000). thus, although limited temporal overlap is possible, these sands are mostly younger than the upper kimmeridgian heno formation deposited on the plateaus to the east of the mads and inge highs. it is likely, therefore, that volgian shoreface sandstones were also deposited in danish territory just east of the mid north sea high along the western flank of the outer rough basin (fig. 1). the shoreface sandstones may have been trapped due to rotation of the outer rough basin during the early–middle volgian. 399 sand deposited by tidal currents? during the kimmeridgian, the marine embayments formed by the tail end and feda grabens were at times separated by land and at other times connected by shallow seas upon the intervening structural highs (fig. 14). during periods of shallow submersion, the sandy plateau between the tail end and feda grabens may have been swept by tidal currents as water masses were transported from one basin to the other over the submarine plateau. the width of the passage over the submarine plateau from the mandal high to the inge high was probably about 35–40 km. it is likely, therefore, that some of the sandstones of the heno formation were deposited or redeposited by tidal currents although investigation of the importance of tidal processes in this setting is precluded by the pervasive bioturbation and the resultant scarcity of primary sedimentary structures. tidal sediments have been reported from broadly equivalent strata in the uk sector, both from the nearby fife and angus fields (spathopoulos et al. 2000) and from the fulmar formation farther north-west (howell et al. 1996). the former authors also described facies in the ‘fife sandstones’ that were interpreted as the deposits of high density flows, either induced by storms or representing turbiditic flows derived from a nearby structural high. acknowledgements the present paper is part of a ph.d. thesis, supported financially by the geological survey of denmark and the ministry of energy (ens no. 1313/92-0002) and supervised by finn surlyk, university of copenhagen. finn surlyk is thanked for his penetrative, constructive comments on the paper and improvement of the english. ron steel and alf ryseth reviewed the paper and are thanked for many helpful comments, which greatly improved the quality of the manuscript. the editorial work by jon r. ineson is very much appreciated. lars henrik nielsen, jon r. ineson, jan andsbjerg and karen dybkjær at geus are thanked for fruitful and stimulating discussions. references andsbjerg, j. 2003: sedimentology and sequence stratigraphy of the bryne and lulu formations, middle jurassic, northern danish central graben. in: ineson, j.r. & surlyk, f. 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(ed.): petroleum geology of northwest europe: proceedings of the 4th conference, 459–468. london: geological society. walker, r.g. & bergman, k.m. 1993: shannon sandstone in wyoming: a shelf-ridge complex reinterpreted as lowstand shoreface deposits. journal of sedimentary petrology 63, 839–851. 402 manuscript received 28 november 1995; revision accepted 15 january 2000. geological survey of denmark and greenland bulletin 33, 2015, 13-16 13 down-hole permeability prediction – a chemometric wire-line log feasibility study from a north sea chalk well kim h. esbensen, niels h. schovsbo and lars kristensen permeability in chalk depends primarily on porosity but also on other factors such as clay and quartz content, and can theoretically be described by the kozeny equation using empirically determined constants (mortensen et al. 1998; røgen & fabricius 2002). recent attempts to predict permeability from wire-line logs have shown that compressional velocity within operative chalk units, defi ned by specifi c surface and hydraulic properties established from stratigraphy and core plugs, can provide excellent well permeability predictions (alam et al. 2011). high-quality predictions depend on a solid knowledge of a multitude of parameters of the relevant ‘operative rock types’. th e more detailed this a priori knowledge is, the better predictions can be achieved. but this approach may, or may not, be fast enough for wellsite operations or when core data are lacking. in this study, we illustrate a situation for direct permeability prediction if only well-site, wire-line logs are available. th is pilot study is based on multivariate descriptor relationships, specifi cally aimed at direct permeability prediction, using all immediately available wire-line characteristics and/or core (plug) information in a top-down mode with sequential exclusion of non-correlated, irrelevant variables. we show prediction-model results based on [log] data only and on [log + plug] data. other relevant descriptors could be included in an augmented x-matrix, such as quantitative core and facies descriptions while still retaining the fast wellsite perspective. however, such data were not included in this feasibility study. material and methods core and log data are from the m-1x well in the danish part of the north sea; core data were collected in the mid1990s during a multi-disciplinary reservoir study (dons et al. 1995). th e m-1x well intersects the danian ekofi sk formation and the maastrichtian tor formation (kristensen et al. 1995). core analysis included determinations of conventional he-porosity and air permeability, whole-rock ca, mg, fe, mn and sr concentrations, δ13c and δ18o isotope ratios, per cent carbonate and per cent non-carbonate. before data analysis all concentrations were corrected to represent weight pr. volume. m-1x was drilled in 1971 on the dan field structure (fig. 1), and encountered a c. 200 m thick hydrocarbon-bearing zone in the chalk. petrophysical evaluation shows the top reservoir is at 1800 m; a gas cap was encountered down to 1880 m and the oil–water contact was found at 2036 m. a 192 m long core was collected from the hydrocarbon-bearing zone with a core recovery of c. 75%. wire-line logs included gamma ray (gr), sonic, formation density, spontaneous potential (sp), calliper, induction log (deep resistivity), lateral log (deep resistivity), micro-lateral log (shallow resistivity; mll) and short normal resistivity (medium resistivity). core data depth and well-log readings were adjusted and aligned applying an estimated common depth shift of 3 m. log readings were sampled for each plug depth to ensure a common plug-log training data set. two chemometric techniques were used, principal component analysis (pca) and partial least squares (pls) regression. pca transforms a matrix of measured data (n samples, p variables), x, into sets of projection sub-spaces delineated by principal components (each a linear combination of all p variables), which display variance-maximised interre© 2015 geus. geological survey of denmark and greenland bulletin 33, 13–16. open access: www.geus.dk/publications/bull 4°e gas field in chalk oil field in chalk inversion zone fault zone studied well 55°30´n dan field m-1x 10 km uk germany norway denmark the netherlands 500 km fig. 1. location of well m-1x in the dan field in the danish part of the north sea. 1414 lationships between samples and variables respectively (martens & næs 1989; höskuldsson 1996; esbensen 2010). pca score plots display groupings, or clusters, between samples based on compositional similarities, as described by the variable correlations (shown in accompanying loading plots), and also quantify the proportion (%) of total data-set variance that can be modelled by each component, see fig. 2. all data analyses in this work are based on auto-scaled data [x-x(avr)/std]. pls regression replaces the classical multiple linear regression and allows direct correlations to be modelled between y and the multivariate x data, among other compensating for debilitating co-linearity between xvariables, (martens & næs 1989; höskuldsson 1996; esbensen 2010). pls regression models are used extensively in science, technology and industry for prediction purposes where the critical success factor is proper validation (esbensen & geladi 2010). both pca and pls result in informative score plots, loading plots (pls: loadingweights) and prediction validation plots, which are the prime vehicles for detailed interpretation of complex data relationships. pls components are based on [x,y] covariance optimisation, but the scientifi c interpretation of the derived scores and loading-weights plots follows procedures which are identical to the pca. validation was based on a test set prepared before modelling: as the m-1x data set is limited, it was sorted with respect to the rhob mg nonc fe dt sp gr mn mll sr il por perm sn ll ca carb ekofisk fm tor fm hod fm δ18o δ13c pca 1 (38%) –8 –4–6 –2 0 2 4 pc a 2 (3 1% ) 6 4 2 0 –2 –4 –6 –8 pca 1 (38%) –0.4 –0.2 0.0 0.2 0.4 0.4 0.2 0.0 –0.2 –0.4 pc a 2 (3 1% ) a b fig. 2. principal component analysis. a: loading and b: score relations for the full training data set (ekofisk, tor and hod formations). the plot models 69% of the total data variance, the proportions are shown along each component axis (38 + 31%). a: abbreviations see fig. 3. –6 –4 –2 0 2 4 6 8 p l s 2 (3 5 % , 2 % ) reference air permeability (md) 0 2 4 6 8 p re d ic te d a ir p er m ea b ili ty ( m d ) pls 1 pls 2 pls 4 pls 6 % ex p la in ed y -v ar ia n ce pls1 (39%, 86%) pls 5 pls 7pls 3 dt gr il ll mll rhob sn sp carb nonc ca mg fe mn sr por perm pls1 (30%, 86%) –0.4 –0.2 0.0 0.2 0.4 0.6 p l s 2 ( 3 5 % , 2 % ) δ18o δ13c y = 0.44 + 0.88x r2 = 0.83 c d ba 4 2 0 –2 –4 –6 –8 0.6 0.4 0.2 0.0 –0.2 –0.4 –0.6 100 80 60 40 20 0 8 6 4 2 0 fig. 3. pls regression model [log + plug] variable set; full training set with ekofisk, tor and hod formations. a: pls x-space score plot (t1-t2). b: corresponding loading-weights plot (w1-w2). c: modelled y-variance. d: prediction versus reference plot. two outliers were deleted from the original data set. proportions of total data variance modelled shown along each pls component [x%, y%]. gr: gamma ray. dt: compressive wave interval travel time. rhob: formation density. il: induction log. ll: lateral log. mll: micro lateral log. sn: short normal resistivity. sp: spontaneous potential. por: he-porosity. perm: air permeability. ca: calcium. mg: magnesium. fe: iron. mn: manganese. sr: strontium. carb: carbonate volume content [calculated]. nonc: non-carbonate [calculated as 100% – carbonate volume %]. for data analysis, concentrations were transformed to weight per volume rock values. legend see fig. 2 15 full permeability range before being randomly split into two independent data sets, i.e. the training versus the test set, securing a realistic prediction performance validation (esbensen 2010; esbensen & geladi 2010). results th ere is a marked and fundamental diff erence in rock properties between the ekofi sk formation and the tor and hod formations (fig. 2). th e ekofi sk formation shows a high concentration of non-carbonate, fe and mn and high gr and mll levels. th ese characteristics are well-known from the north sea region, which forces a cautious approach to data set defi nition. th e developed permeability model may, or may not, apply to both the tor and the hod formations and the ekofi sk formation. th is will depend on whether the relationships between the x data from the three formations are similar with respect to correlation to permeability. a two-component pls model on the full (log + plug) variable set predicts permeability with satisfactory validation results as seen in the prediction versus reference plot in fig. 3 (slope 0.88; r2 = 0.83), suggesting that the pls model leads to better permeability estimates than normally achieved from conventional poro-perm plots. conventional statistics pertaining to a fi tted linear regression model between predicted (y) versus reference (x) values are used to express the degree of prediction strength: slope and regression coeffi cient, r2. for both these modelling indices the criterion is to be as close to 1.00 as possible. such validation statistics must be based on proper validation (esbensen & geladi 2010). th e permeability model is primarily carried by positively correlated por, ll, il, sn and negatively correlated rhob and gr, but several other log and composition variables also have minor, but signifi cant infl uence. from the loading-weights plot it is diffi cult to resolve any fully irrelevant variables; pls models benefi t from using a full x-variable complement; variable selection is not needed in this case. variable relationships are interpreted in the more appropriate pls loading-weight plots; a technical detail not to be elaborated on here, as interpretation follows the same principles (martens & næs 1989; esbensen 2010). figure 4 shows permeability prediction only based on log data (ekofi sk formation excluded), simulating a situation in which there are only well-site, wire-line logs available for the fastest possible permeability prediction. th e validation results for this model (slope 0.77; r2 = 0.75) are lower, but still acceptable for direct on-site permeability screening based on fig. 4. pls regression model (logs only). a: pls x-space loading weights plot (t1-t2). b: prediction versus reference plot. proportions of total data variance modelled shown along each pls component [x%, y%]. legend see fig. 2, abbreviations see fig. 3. –0.6 –0.4 –0.2 0.0 0.2 0.4 0.6 pls 1 (71%, 84%) il 0.4 0.2 0.0 –0.2 –0.4 –0.6 –0.8 6 4 2 0 p re d ic te d a ir p er m ea b ili ty ( m d ) p l s 2 ( 1 3 % , 4 % ) 0 2 4 6 8 reference air permeability (md) y = 0.61 + 0.77x a b rhob dt sp mll ll perm sn gr r2 = 0.75 fig. 5. reservoir properties versus depth. a: wire-line density log and core porosity measurements. b: predicted air permeability based on the model presented in fig. 4, compared with reference permeability (core measurements). the reservoir is gas filled from 1800 to 1880 m and oil filled down to a depth of 2036 m. the permeability model does not apply to the ekofisk formation (red rectangle). 1800 1850 1900 1950 2000 2050 tor hod ekofisk a b 2.4 2.0 0 5 10 4020 d ep th (m ), m d core porosity (%) permeability predictor (x=logs, tor and hod), md density (g/cm3) core measurement 1616 contemporaneous log data alone. th e results in figs 3 and 4 indicate that the tor formation can be modelled equally well with, or without, the hod formation. figure 5 shows stratigraphic permeability results for the all-logs prediction model (fig. 4), plotted together with measured core porosity (%) and density. an all-logs prediction model is fully able to characterise the hod and tor formations, but not the ekofi sk formation. for the latter, additional core information is necessary (fig. 3). discussion th e compositional diff erence between the ekofi sk formation and the tor formation has also previously been studied by multivariate data analysis (kunzendorf & sørensen 1989), pointing to a relationship between reservoir quality and geochemistry. røgen & fabricius (2002) showed that these compositional and textural relations are also refl ected in specifi c surface area diff erences between the formations, and thus in permeability and porosity diff erences. our analysis shows that high permeability is closely related to high porosity, and to high resistivity (fig. 3; ll, il, sn), whereas low permeability is related to high density and high gr, high non-carbonate content and thus to impure chalk with high concentrations of mn, fe and mg. røgen & fabricius (2002) also showed that quantitative mineral data can help to explain permeability values better. our analysis also shows that permeability predictions from wire-line logs alone strongly depend on the sonic and resistivity logs (fig. 4; dt, il, ll, sn and sp). th ese fi ndings complement those of alam et al. (2011) in which permeability was also predicted but based on the sonic log alone (dt). our analysis further shows that it is pos sible to model permeability more comprehensively by including the full set of readily available wire-line logs. conclusions th e present study confi rms that multiple parameters control permeability levels. both log data and core data can be used advantageously in direct pls prediction; there are real benefi ts in including the full set of available well-site parameters. prediction of permeability from models based on log information alone is useful for screening purposes, whereas permeability prediction from models based on both log data and core data are, not surprisingly, signifi cantly better. which approach to use depends on the context in which permeability prediction is used, especially on the time available for securing the additional core information from the laboratory. th is study shows that direct well-site permeability prediction is feasible. improvements can be made by adding standard he-porosity data and other easily measured conventional laboratory core parameters. th e feasibility study was based on a 192 m long chalk interval in a single well only. th e database can be extended to include more of the comprehensive core data available from the danish north sea. based on an augmented data set, it is in principle an easy task to refi ne this pilot study to investigate the more general limits of the feasibility demonstrated. a parallel study based on a similar approach using log data and logs + core data also proved successful for prediction of ‘functional rock types’ for other lithologies than chalk, i.e. alum shale (schovsbo et al. 2015). functional rock types may correlate with rock strength and can here be used for optimisation of the completion design. references alam, m.m., fabricius, i.l. & prasad, m. 2011: permeability prediction in chalks. aapg bulletin 95, 1991–2014. dons, t., jacobsen, f. & stentoft , n. 1995: chalk diagenesis and reservoir properties – dan fi eld case study. dgu service report 15, 2 vols, 251 pp., 259 pp. københavn: danmarks geologiske undersøgelse. esbensen, k.h. 2010: multivariate data analysis, in practise. an introduction to multivariate data analysis and experimental design, 5th edition, 598 pp. oslo: camo. esbensen, k.h. & geladi, p. 2010: principles of proper validation: use and abuse of re-sampling for validation. journal of chemometrics 24, 168–187. höskuldsson, a. 1996: prediction methods in science and technology 1, basic theory, 405 pp. holte: th or publishing. kristensen, l., dons, t., maver, k.g. & schiøler, p. 1995: a multidisciplinary approach to reservoir subdivision of the maastrichtian chalk in the dan fi eld, danish north sea. aapg bulletin 79, 1650–1660. kunzendorf, h. & sørensen, p. 1989: geochemical criteria for reservoir quality variations in chalk from the north sea, 99 pp. roskilde: risø national laboratory. martens, h. & næs, t. 1989: multivariate calibration, 419 pp. chichester: wiley. mortensen, j., engström, f. & lind, i. 1998: th e relationship among porosity, permeability, and specifi c surface of chalk from the gorm field, danish north sea. spe reservoir evaluation & engineering 13, 245–251. røgen, b. & fabricius, i.l. 2002: infl uence of clay and silica on permeability and capillary entry pressure of chalk reservoirs in the north sea. petroleum geoscience 8, 287–293. schovsbo, n.h., esbensen, k.h., nielsen, a.t., derbez, e., gaucher, e.c., poirier-coutansais, x., riou, a., tallone, p. & milton-taylor, d. 2015: rock types in the scandinavian alum shale resource play: defi nitions and predictions. 77th eage conference & exhibition, madrid, 1–4 june, 2015. abstract. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ke@geus.dk geological survey of denmark and greenland bulletin 41, 2018, 67-70 67 the european trading and whaling activities of the 17th– 19th centuries provide records of climate and seaice conditions off west greenland in the form of ships’ logs and other official documents in many archives around europe. these documents, combined with evidence from marine sediments, help describe climate changes in general, and seaice volume changes in particular, in connection with human activity in the region. the greenland national museum & archives in nuuk (nka) hosts a unique collection of original documents presenting detailed insight into weather and ice conditions as well as the daily life of the colonial centres and outposts recorded by the documents of the danish administration. these documents also reveal many aspects of the interaction between the inuit and europeans from 1779 onwards. information retrieved from the archives in nuuk has been combined with results from palaeo-environmental investigations of marine sediment cores to unravel climate variability and changes in sea ice. this information has been supplemented with data from an extensive field programme using drones to document onshore remains from the whaling period in the disko bugt region (fig. 1). history of european whaling off west greenland european arctic whaling around svalbard and jan mayen declined in the early 18th century, and it has been debated whether over-exploitation (e.g. nansen 1924) or climate change (e.g. vibe 1967) was responsible for the decline. european summer temperature records (luterbacher et al. 2016) indicate that the decline coincided with a marked climate cooling in europe. high hunting pressure on the whales and thick sea-ice cover in svalbard waters early in the 18th century apparently forced the whale population to seek away from inshore areas into the ice-loaded waters farther offshore. as a consequence, european whaling companies began to look for other areas, which led them to explore the whale resources of the davis strait. taking a general ‘seesaw’ winter-climate pattern between (west) greenland and europe into consideration (seidenkrantz et al. 2008), in contrast to the cold conditions around svalbard (luterbacher et al. 2016), the sea-ice and whaling conditions in disko bugt around 1700 were likely more favourable. a large part of the european whaling fleet therefore moved their activities to the davis strait off west greenland where the disko bugt region became the focal point of whaling activities. in the region, the european whalers’ contact with the local inuit population significantly impacted their culture and living conditions through the exchange of goods and social interaction. in 1719, an increased number of dutch whaling vessels started to operate in the davis strait (leinenga 1995; hacquebord 2006) with the greenland (bowhead – balaena mysticetus) whale as the main target. whaling activities in the european trading, whaling and climate history of west greenland documented by historical records, drones and marine sediments naja mikkelsen, antoon kuijpers, sofia ribeiro, mikkel myrup, inge seiding and ann e. lennert disko disko bugt qeqertarsuaq qasigiannguit ilulissat por-16 aasiaatpullat greenland ice sheet rodebay hunde ejland g re en la nd 69˚ 55˚ 53˚ 51˚ 69˚ fig 1. disko bugt was the focus of whaling activities in past centuries. areas investigated in 2016 and 2017 were rodebay embayment, hunde ejland and pullat. a sediment core por16_rb#1 was retrieved just outside the rodebay embayment. fig 2. whales caught by danish shore-based whaling stations were numbered and often drawn as cartoons in danish reports. this document clearly illustrates that whaling had been successful for the period reported (source nka archives, nuuk). © 2018 geus. geological survey of denmark and greenland bulletin 41, 67–70. open access: www.geus.dk/bulletin http://www.geus.dk/bulletin 6868 davis strait were concentrated around disko bugt. in 1721, denmark affirmed sovereignty over greenland and in 1776 established the state-owned royal greenland trade department (kgh) with whales and seals as the primary commodities. this danish initiative had difficulty competing with other highly efficient whaling fleets and after disappointing catches, further attempts with danish sea-going whaling ceased after 1780. instead danish activities primarily used shore-based stations for sealing and whaling and this hunting turned out to be economically successful in the following decades. after the napoleonic wars, european demand for whale products increased, and new danish coastal whaling projects were started in the davis strait area (e.g. the island of pullat off the coast at aasiaat – egedesminde). however, due to increasing industrial use of coal and other types of oil in europe and as a consequence of a series of severe ship losses early in the 19th century, the davis strait whaling lost its importance. evidence from historical archive data data from several whaling settlements are available in the archives in nuuk and provide a wealth of detailed information after 1779 on weather, sea ice occurrence, number of whales caught (fig. 2) as well as results from hunting, fishing, travelling, social life, health conditions, religious celebrations and occasional visits by european whalers (nka archives). successful coastal whaling was only possible with participation of the local inuit population, in many cases also involving women. the european whaling activities thus had significant impact on the previously prosperous inuit hunting and their traditional way of life. the inuit women’s national costume rich in bead decoration, and various forms of folk dance music are examples of social interaction between the two groups. although dutch whaling activities along the west greenland coast had virtually ceased by 1800, a series of geographical dutch names are a reminder of this period, such as: vaigat (dutch for windy passage), rodebay, fortunebay, grønne ejland, and hunde ejlande. in the early 20th century, scottish whalers, the only group still active in the area, moved their centre of activity northwards and made port calls at upernavik instead of disko bugt. as can be seen in the historical records from upernavik (fig. 3), this period is clearly characterised by longer seaice seasons than in the later part of the 19th century suggesting that better whaling conditions in the early 20th century had shifted northwards into more ice-affected areas along the west coast of greenland. tracing past whaling activities onshore the coastal areas of the disko bugt region display a rich variety of physical remains from the whaling period. to locate and investigate former whaling sites, mikkel myrup from fig. 3. dates of freezing and ice breakup in the upernavik district plotted from archival data. the time of sea-ice breakup shows a clear shift, with the sea-ice duration increasing by c. 40 days over a period of c.  30 years (c. 1890–1920; a. kuijpers, unpublished data 2014). fig. 4. remains from the coastal whaling period on the island of pullat are documented on the orthophotomosaic based on drone flights. a: winter house. b: the main structures from the whaling station period 1823–1825. c: winter houses. d, e: grave fields. photo: m. myrup 2017. 1790 10. jul 30. jun 20. jun 10. jun 31. may 21. may 11. may 01. may 11. dec 16. dec 06. dec 01. dec 25. nov 21. nov 15. nov 11. nov 06. nov 1810 1830 1850 1870 1890 1910 1930 1790 1810 1830 1850 1870 1890 1910 1930 start of fast sea ice start of ice break-up 68.730˚n 68.726˚n 68.722˚n © mikkel myrup pullat 0 m 200 53.120˚w 53.110˚w 69 nka in 2017 undertook a number of drone operations in three areas in the disko bugt region: oqaatsut (rodebay), kitsissuarsuit (hunde ejlande) and pullat (fig. 1). the small greenlandic settlement oqaatsut (rodebay) is located in a protected embayment just north of ilulissat. in the 1600s, this area was called ‘roo baj – red bay’ by the dutch. whales caught locally by traditional whaling techniques were pulled ashore on the flat rocks where blood from the butchered animals coloured the small embayment red. today remains from the whaling activities are visible in the form of natural butchering places created by rock depressions with winches still present. amongst other structures still visible are the remains of a whaling station established by danish authorities on the island of pullat in 1823 (fig. 4). it was a sudden and intense effort by kgh to boost the coastal-based whaling activity in the area and also to counteract the strong presence of european whalers interacting with the inuit farther north. the whaling station at pullat, however, was closed after only two years due to unsuccessful whale hunting. towards the middle of the 19th century, kgh eventually gave up the whaling outposts in the disko bugt region and focused its efforts on seal hunting and procuring other greenlandic products. the graveyard on kitsissuarsuit (hunde ejlande) hosts modern graves with traditional white-painted wooden crosses and a central section from the european whaling period. this central section known as ‘the old whalers graveyard’ is occupied by c. 100 graves covered by elongated stone piles. unlike the modern white crosses, two of these graves are marked with plain wooden, squared posts (fig. 5). similar old graves are reported from the graveyard ‘liknesset’ on north-western svalbard, which was established by dutch whalers in the 17th century. evidence from marine sediment records in august 2016, a sediment coring campaign was undertaken from the research vessel porsild. twelve sediment cores were retrieved from the rodebay area to obtain a record of climate and hydrographic changes and possible environmental effects of human activity in this area spanning the whaling period. cores collected from the enclosed inner part of the rodebay area consist of unconsolidated mud with high organic contents are underway. dna analyses will be performed on these cores to identify which mammalian species were slaughtered at the site throughout the whaling period. to capture changes in the marine environment a core (core por16_rb#1) was retrieved just outside the rodebay embayment. the geochronology of this 158 cm core was based on 210pb and 137cs analyses of the upper 15 cm, and 3 ams 14c dates on marine bivalve shells. the age-model was constructed using bayesian analyses with the bacon package available in the r platform. the radiocarbon offset used was 140 ± 60 years, and radiocarbon dates were calibrated using the marine13 calibration curve. the core was xrf scanned and analysed in terms of magnetic susceptibility and total organic carbon (toc) content (fig. 6). the prominent dark colouration of the interval corresponding to the period c. ad 1600–1750 reveals the presence of poorly oxygenated sediments. locally enhanced organic matter fluxes can be expected at this site to originate from whale butchering colouring the sediments black. however, this dark interval precedes by c. 100 years the beginning of intensive whaling activity in the region, and the organic carbon content does fig. 5. among the graves in the ‘old whalers graveyard’ on kitsissuarsuit (hunde ejlande), two graves are marked with plain squared timbers. this type of grave marking has only been found in the old whalers’ graveyard on svalbard established in the 17th century. modern graves with the traditional white painted wooden crosses are seen in the background. 0 100 200 300 400 500 600 700 800 900 1000 1100 1200 1300 1400 1500 c or e de pt h (m m ) 0.6 0.9 1.2 toc (%) 160 360 k 240 440 640 ca 0 100 br 30 60 90 magnetic susceptibility 1300 1400 1500 1600 1700 1800 1900 2000 a ge ( a d ) fig. 6. sediment core por16_rb#1 was retrieved just outside the rodebay embayment. the geochronology of this 158 cm core was based on 210pb and 137cs analyses of the upper 15 cm, and 3 ams 14c dates on marine bivalves. the prominent dark colouring of the interval corresponding to the period c. ad 1600–1750 of the little ice age reveals the presence of non-oxic conditions most likely due to persistent sea-ice cover. sediments deposited after ad 1600 display regular oscillations in k and ca with a duration of close to 80 years. 7070 not indicate enhanced fluxes in organic matter. instead, local hydrographic conditions with reduced bottom-water oxygenation are likely to be responsible for the dark layers. this sediment interval corresponds to one of the more severe little ice age cooling episodes during which the sea-ice cover along the coast of the inner part of disko bugt may have persisted throughout most of the year (ribeiro et al. 2011). during cold intervals of the little ice age, the mean annual temperature in the disko bugt region area was 2–4°c lower than present day values (humlum 1999). prior to this period of generally colder climate, about 1000 years ago, the regional atmospheric climate was relatively mild in west greenland, but soon after, i.e. close to ad 1200, average temperatures started abruptly to decline by as much as 4°c over about 80 years (d’andrea et al. 2011). this climate deterioration involved an increase of widespread ice formation in the inshore waters and fjords of west greenland (kuijpers et al. 2014). regional climate since then has remained generally colder until about 150 years ago, although with large fluctuations (ribeiro et al. 2011). the xrf data of core por16_rb#1 reveal regular oscillations in the upper part of the core that appear to occur on a multi-decadal time scale. during the past 1000 years, 60–80 years periodicity is noted in sea-surface conditions of the disko bugt areas (allan et al. 2018) suggesting clear linkages with the ‘atlantic multidecadal oscillation’ (amo) in which warmer and more saline north atlantic surface-water conditions alternate with cooler and less saline conditions around ad 1910 (reverdin 2010). based on historical information from the kgh records from upernavik (a. kuijpers, unpublished data 2014), the temperature and salinity may also be reflected by the duration of the sea-ice season off west greenland, which in the beginning of the 20th century extended into june (fig. 3). such a relationship between winter sea-ice conditions in the north-western north atlantic and ‘atlantic multidecadal oscillation’ has also been reported by chan et al. (2017), who found that during the past four centuries an extended sea-ice cover and associated reduced ice productivity can be linked to the cool phase of the ‘atlantic multidecadal oscillation’. acknowledgements this study was part of the eu funded project ice-arc (ice, climate and economics in the arctic) and supported by funding from the european union 7th framework programme, grant number 603887 and by funding from the national museum and archives (nka), nuuk, project encounters. niels frandsen, former head of the archive, is gratefully acknowledged for research support (ak) in nuuk in 2014. references allan, e., de vernal, a.e., knudsen, m.f., hillaire-marcel, c., moros, m., ribeiro, s., oulet-bernier, m.-m. & seidenkrantz, m.-s. 2018: late holocene sea surface instabilities in the disko bugt area, west greenland, in phase with δ18o oscillations at camp century. paleoceanography and paleoclimatology 33, 227–243, http://dx.doi. org/10.1002/2017pa003289 chan, p., halfar, j., adey, w., hetzinger, s., zack, t., moore, g.w.k., wortmann, u.g., williams, b. & hou, a. 2017: multicentennial record of labrador sea primary productivity and sea-ice variability archived in coralline algal barium. nature communications 8, 15543, 10 pp., http://dx.doi.org/10.1038/ncomms15543   d’andrea, w.j., huang,y., fritz, s.c. & anderson, n.j. 2011: abrupt holocene climate change as an important factor for human migration in west greenland. proceedings of the national academy of sciences of the united states of america 108(24), 9765–9769. hacquebord, l. 2006. two centuries of bowhead whaling around spitsbergen; its impact on the arctic avifauna. in: ringstad, j.e. (ed.): whaling and history ii, new perspectives, 87–94. sandefjord: kommandør chr. christensens hvalfangst museum. humlum, o. 1999: late-holocene climate in central west greenland: meteorological data and rock-glacier isotope evidence. the holocene 9(4), 581–594. kuijpers, a., mikkelsen, n., ribeiro, s. & seidenkrantz, m.-s. 2014: impact of medieval fjord hydrography and climate on the western and eastern settlements in norse greenland. journal of the north atlantic, special volume 6, 1–13. leinenga, j.r. 1995: arctische walvisvangst in de achttiende eeuw, de betekenis van straat davis als vangstgebied, 237 pp. phd thesis, university of groningen. amsterdam: de bataafsche leeuw. luterbacher, j., werner, j.p. & smerdon, j.e. 2016: european summer temperatures since roman times. environmental research letters 11(2), 024001. http://dx.doi.org/10.1088/1748-9326/11/2/024001 nka archives. inspektoratet for nordgrønland. modtagne embedsdagbøger & journaler: arkiv nr. 01.02 o7.20 10; nr. 01.02. 07.20. 14; nr. 01. o2. 07.20. 15; nr. 01.02. 19.20. 3: nr. 01. o2 73.45. 3. nka 01.01/13.22/: 01.02/73.42/7; 01.02/13.22/9. nansen, f. 1924: blant sel og bjørn. min første ishavs-ferd, 285 pp. kristiania: jacob dybwads forlag. reverdin, g. 2010: north atlantic subpolar gyre variability (1895– 2009). journal of climate 23, 4571–4584. ribeiro, s., moros, m., ellegaard, m. & kuijpers, a. 2012: climate variability in west greenland during the last 1500 years – evidence from a highresolution marine palynological record from disko bay. boreas 41, 68–83. seidenkrantz, m.-s., roncaglia, l., fischel, a., heilmann-clausen, c., kuijpers, a. & moros, m. 2008: variable north atlantic seesaw patterns documented by a late holocene marine record from disko bay, west greenland. marine micropaleontology 68, 66–83. vibe, c. 1967: arctic animals in relation to climatic fluctuations. meddelelser om grønland 170(5), 227 pp. authors’ addresses n.m., a.k. & s.f., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: nm@geus.dk. m.m. & i.s.greenland national museum & archives, hans egedesvej 8, 3900 nuuk, greenland. a.e.l.,the polar museum, søndre tollbodgate 11b, 9008 trømsø, norway. http://dx.doi.org/10.1002/2017pa003289 http://dx.doi.org/10.1002/2017pa003289 http://dx.doi.org/10.1038/ncomms15543 http://dx.doi.org/10.1088/1748-9326/11/2/024001 mailto:nm@geus.dk geological survey of denmark and greenland bulletin 31, 2014, 87-90 87 mass loss from an ice-sheet drainage basin in west greenland morten l. andersen, signe b. andersen, lars stenseng, henriette skourup, william colgan, steen s. kristensen, john p. merryman boncori, andreas p. ahlstrøm, xavier fettweis, rené forsberg, michele citterio, jason e. box, dirk van as and robert s. fausto the greenland ice sheet is losing mass to the ocean at an increasing rate (thomas et al. 2006). during the 1980s the ice sheet was believed to be in near-equilibrium (van den broeke et al. 2009). within the first decade of the 21st century, however, a net negative balance was observed. greenland’s present rate of ice loss is c. 250 gt yr–1, equivalent to a sea-level rise contribution of c. 0.69 mm yr–1. the rate of ice loss has increased over the post 1992 observation period (shepherd et al. 2012). the ice-sheet mass budget can be partitioned into two main components: (1) surface mass balance (smb; the net difference between accumulation and surface ablation) and (2) marine ice loss (d; iceberg discharge via glacier dynamics plus subsurface melt at the glacier terminus). over the past decade, the surface mass-balance proportion has accelerated relative to the d component, changing from c. 50% in 2000–2008 (van den broeke et al. 2009) to more than two thirds (68%) in 2009–2012 (enderlin et al. 2014). whereas modern climate models appear to capture the surface mass-balance response to climate change, the physical processes driving variability in glacier discharge are more complex. recent increases in d may be due to: (1) changing force-balance at the ice-ocean interface as suggested by model simulations (e.g. nick et al. 2009, 2013), (2) changing basal lubrication at the ice-bed interface due to increased meltwater availability (zwally et al. 2002; andersen et al. 2010), and/or (3) decreasing ice viscosity due to increasing ice temperature (van der veen et al. 2011). the high spatial variability in these forcing mechanisms and a large sensitivity to local fjord geometry (nick et al. 2013) require basin scale studies of glacier dynamics to elucidate local causes of glacier acceleration. in 2007 the programme for monitoring the greenland ice sheet (promice) was initiated to gain insight into the changing mass balance of the greenland ice sheet using quantitative meteorological observations, as well as airborne surveys of ice thickness and flow-velocity observations (ahlstrøm et al. 2008). here we present the first calculations of ice discharge using promice observations, with focus on a west greenland ice-sheet drainage basin previously defined as ‘basin 7’ (zwally et al. 2012; fig. 1). the c. 400 km long ice-sheet margin within basin 7 includes the 6 km wide jakobshavn isbræ, and several other marine-terminating outlet glaciers, such as store gletscher and rink isbræ (figs 1, 2). we combine satellite-derived, ice-surface velocities, airborne ice-thickness measurements, and modelled surface mass balance to assess the dynamic discharge from basin 7. © 2014 geus. geological survey of denmark and greenland bulletin 31, 87–90. open access: www.geus.dk/publications/bull studied basin area ends of section in fig. 2 2007 flight line 2011 flight line computed surface velocities 250 km 60°n 50°w 60° 50° 40° 30° 20° 10° 0° 40°w 80°n 70°n rink isbræ store gletscher jakobshavn isbræ basin 7 fig. 1. map of greenland showing interpolated flight lines for 2007 and 2011, area of computed surface velocities for this study (fig. 3) and the studied catchment area. l 8888 data and methods we estimate the solid ice discharge (d) into the ocean according to the input–output method of rignot & kanagaratnam (2006). first we quantify the mass flux (f) discharging across a flux gate, upstream of the boundary between the ice sheet and the ocean (the grounding line), defined by the path of the promice airborne ice-thickness surveys conducted in the summers 2007 and 2011 (fig. 1). the elevation of this flux gate is c. 1500 m a.s.l. in basin 7 (figs 2, 3). the ice-surface and bed elevations determined by the airborne surveys were interpolated to c. 30 m spacing along the flux gate to resolve spatial variability in ice flow. the flux f at grid point i is computed as fi = hi ∙ li ∙ vi, where hi is the ice thickness, li is the spacing along the flight line (c. 30 m), and vi is the depth-averaged ice velocity component that is perpendicular to the flux gate. ice surface velocities were derived by applying offset tracking to alos/ palsar synthetic aperture radar (sar) data acquired between november 2009 and february 2010, using the susie processing chain based on the commercial package gamma (merryman boncori et al. 2010; ahlstrøm et al. 2011). uncertainties associated with the ice velocities were estimated using the method of mohr & merryman boncori (2008) and are under 10% (fig. 3). we assume a uniform vertical velocity profile, where ice-surface velocity is equivalent to depth-averaged velocity (i.e. ‘plug flow’; rignot & kanagaratnam 2006). with the total basin flux f (=∑fi) known, the grounding line discharge (d) can be estimated by adding the spatially integrated surface mass balance (smb) of the area downstream (‘ds’) of the flux gate: d = f + smbds,ref, where smbds,ref is a reference period (1961–1990) mean smb field from the regional climate model mar v3.2, forced at its boundaries by ecmwf reanalysis data and run at a spatial resolution of 25 km (fettweis et al. 2013a). similarly, the mass balance upstream of the flux gate (interior mass balance, imb) can be computed by subtracting f from the upstream spatially integrated smb for the reference period: imb = smbus,ref –f. quantification of d allows us to estimate the total mass balance (tmb) of the drainage basin. the tmb value is calculated as tmb = smbtot,yr –d, where smbtot,yr is the yearly smb spatially integrated across the entire basin. estimated uncertainty (σb) on radar-derived bed elevation b values is 80 m and estimated uncertainty (σs) on laser-derived surface elevation observations (s) is 0.1 m. assuming errors in b and s are random, we take uncertainty in ice thickness (σh) as the sum in quadrature of the fractional uncertainties of σb and σs (e.g. colgan et al. 2008). uncertainty in flux f at gridpoint i (σfi) is similarly taken as the sum in quadrature of the fractional uncertainties of σhi and σvi, where the latter term is the uncertainty in the annual fig. 3. surface velocities derived from synthetic aperture radar (sar) data for 2009–2010 used in this study. the contour lines are based on the digital elevation model of the greenland ice mapping project (howat et al. 2014). 0 50 100 150 200 250 300 350 400 –1000 –500 0 500 1000 1500 2000 distance (km) el ev at io n (m a bo ve o r b el ow w g s8 4 el lip so id ) 2007/2011 bed 2007 observed surfacetrough below jakobshavn isbræ south north fig. 2. south-to-north section of the studied basin along the 2007 and 2011 flight lines. for location see fig. 1. 0 500 1000 1500 2000 2500 3000 su rfa ce v el oc ity (m y ea r–1 ) 1500 1500 15 00 69°n 73°n 52°w 2007 flight line 2011 flight line jakobshavn isbræ store gletscher rink isbræ 50 km 89 depth-averaged velocity at i. we assume no uncertainty in li. uncertainty in the total flux f in the basin is then σf = ∑σfi. we take uncertainty in smb to be 15% at basin scale (fettweis et al.  2013a) and similarly propagate uncertainties in both f and smb as the sum in quadrature of fractional uncertainties when assessing the cumulative uncertainty (σd) associated with the grounding-line ice discharge. uncertainties on tmb and imb are developed analogously. the thickness observations were carried out in summer, and we do not account for the difference between summer and winter ice velocities. however, at c. 1500 m a.s.l., we expect the difference between summer and winter ice velocities to be small (<2%; joughin et al. 2008). results both the upstream flux (f) and downstream discharge (d) in basin 7 are within the uncertainty of their respective values in 2007 and 2011 (table 1). whereas we employ different airborne-derived, ice-geometry data for each year, the velocity field used is identical for the two years (winter 2009/2010 values), as is the surface mass-balance correction (1961–1990 values). the similar mass fluxes indicate that changes in ice geometry along the c. 1500 m contour were slight between 2007 and 2011. interior mass-balance values for both years are zero within the uncertainty, which is in good agreement with zwally et al. (2011), who found a slight mass gain of 8 gt yr–1 above 2000 m a.s.l. the total mass-balance values are also, within uncertainty, similar for 2007 and 2011. considering the c. 10 gt yr–1 decrease in d, this suggests that yearly fluctuations in the dynamics of major tidewater outlet glaciers in basin 7 are balanced by variations in surface mass balance. the mean total mass-balance value (–30.5 gt yr–1) corresponds to a sea-level rise contribution of c. 0.08 mm yr–1, and agrees within uncertainty with a satellite gravimetry-derived total mass-balance estimate of –24 ± 1 gt yr–1 for basin 7 over the 2004 to 2010 period (colgan et al. 2014), and a 2007 total mass-balance value reported in rignot et al. (2008) of –36.7 gt yr–1 for an analogous west greenland basin. the 2007 value we present is more negative than a basin 7 estimate of –14 ± 1 gt yr–1 over the 2003 to 2007 period derived from satellite altimetry (zwally et al. 2011). this latter study, however, preceded the 2007 to 2011 observation period, and may therefore reflect the less negative surface mass-balance regime prior to the observation period (fettweis et al. 2013b). summary remarks rignot & kanagaratnam (2006) invoked an assumption of negligible changes in ice geometry between their flux gates and the grounding line. as their flux gates are located at the c. 1000 m elevation contour, any dynamic thickening or thinning signals affect a relatively small proportion of the basin area. given that the promice flux gates are substantially farther inland from the grounding line, we are exploring approaches for explicitly correcting d values for recent changes in ice geometry between the upstream flux gate and the downstream grounding line. this may be particularly relevant in highly dynamic areas, such as the jakobshavn isbræ area. a preliminary assessment of such a correction for basin 7 suggests that the rate of change in downstream ice volume is equivalent to c. 25% of d, which would further decrease the total mass balance by up to 15 gt yr–1. as more synthetic aperture radar data become available, we will improve the temporal coverage of the promice ice-surface velocity product to annual resolution. the plug-flow assumption adds a negative bias to the massloss estimates by assuming that all flow is caused by sliding at the bed, i.e., the surface speed is equal to the mean flow velocity of the ice column. this may be valid in the fast flowing coastal areas, but higher up on the ice sheet the assumption is less valid, where the surface velocity is a mix of sliding and deformation, and the vertically averaged flow speed can be as low as 80% of the observed surface speed. in the promice framework, this basin-scale mass-balance assessment will be extended to deliver basin-scale massbalance and ice-discharge estimates of the entire greenland ice sheet over multiple observation years. this survey aims to improve portioning of mass loss at basin scale, contributing to improved sea-level rise projections for the greenland ice sheet. acknowledgments promice is funded by the danish ministry of climate, energy and building, and is operated by the geological survey of denmark and greenland. this paper is contribution number 38 of the nordic centre of excellence svali, stability and variations of arctic land ice, funded by the nordic top-level research initiative. 2007 79.5 ± 6.1 70.4 ± 6.2 –5.6 ± 12.6 –31.3 ± 8.6 2011 69.7 ± 5.3 60.6 ± 5.5 –4.1 ± 12.3 –29.7 ± 7.2 year upstream ice interior total mass flux (f) discharge (d) balance (imb) balance* table 1. mass fluxes in gt per year * total mass balance = smbtot,yr – d 9090 references ahlstrøm, a. & the promice project team 2008: a new programme for monitoring the greenland ice sheet mass loss. geological survey of denmark and greenland bulletin 15, 61–64. ahlstrøm, a. et al. 2011: promice 2007–2010. final report for the establishment phase of the programme for monitoring of the greenland 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greenland ice sheet surface mass balance contribution to future sea level rise using the regional atmospheric climate model mar. the cryosphere 7, 469–489. fettweis, x., hanna, e., lang, c., belleflamme, a., erpicum, m. & gallée, h. 2013b: brief communication: important role of the mid-tropospheric atmospheric circulation in the recent surface melt increase over the greenland ice sheet. the cryosphere 7, 241–248. howat, i.m., negrete, a. & smith, b. 2014: the greenland ice mapping project (gimp) land classification and surface elevation datasets. the cryosphere discussions 8, 453–478. merryman boncori, j.p., dall, j., ahlstrøm, a.p & andersen, s.b. 2010: validation and operational measurements with susie – a sar icemotion processing chain developed within promice (programme for the monitoring of the greenland ice sheet). in: lacoste-francis, h. (ed.): proceedings of the esa living planet symposium, bergen, norway (esa sp-686) noordwijk: esa communications estec. mohr, j.j. & merryman boncori, j.p. 2008: an error prediction framework for interferometric sar data. ieee transactions on geoscience and remote sensing 46, 1600–1613. nick, f.m., vieli, a., howat, i.m. & joughin, i. 2009: large-scale changes in greenland outlet glacier dynamics triggered at the terminus. nature geoscience 2, 110–114. nick, f.m., vieli, a., andersen, m.l., joughin, i., payne, a.j., edwards, t.l., pattyn, f. & van de wal, r.s.w. 2013: future sea-level rise from greenland’s main outlet glaciers in a warming climate. nature 497, 235–238. rignot, e. & kanagaratnam, p. 2006: changes in the velocity structure of the greenland ice sheet. science 311, 986–990. rignot, e., box, j.e., burgess, e. & hanna, e. 2008: mass balance of the greenland ice sheet from 1958 to 2007. geophysical research letters 35, l20502. shepherd, a. et al. 2012: a reconciled estimate of ice-sheet mass balance. science 338, 1183–1189. thomas, r., frederick, e., krabill, w., manizade,s. & martin, c. 2006: progressive increase in ice loss from greenland. geophysical research letters 33, l10503, doi:10.1029/2006gl026075. van den broeke, m., bamber, j., ettema, j., rignot, e., schrama, e. & van den berg, w.j., van meijgaard, e., velicogna, i. & wouters, b. 2009: partitioning recent greenland mass loss. science 326, 984–986. van der veen, c.j., plummer, j.c. & stearns, l.a. 2011: controls on the recent speed-up of jakobshavn isbræ, west greenland. journal of glaciology 57, 770–782. zwally, h., abdalati, w., herring, t., larson, k., saba, j. & steffen, k. 2002: surface melt-induced acceleration of greenland ice-sheet flow. science 297, 218–222. zwally, h., giovinetto, m., beckley, m. & saba, j. 2012: antarctic and greenland drainage systems, nasa, goddard space flight center, cryospheric sciences laboratory. http://icesat4.gsfc.nasa.gov/cryo_ data/ant_grn_drainage_systems.php. zwally, h. j. et al. 2011: greenland ice sheet mass balance: distribution of increased mass loss with climate warming; 2003–07 versus 1992– 2002. journal of glaciology 57, 88–102. authors’ addresses m.l.a., s.b.a., w.c., a.p.a., m.c., j.e.b., d.v.a. & r.s.f., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: mola@geus.dk l.s., h.s., s.s.k. & r.f., dtu space, dk-2800 lyngby, denmark. x.f., department of geography, university of liège, 4000 liège, belgium. j.p.m.b., istituto nazionale di geofisica e vulcanologia, 00142 rome, italy. geological survey of denmark and greenland bulletin 1, 403-436 403 one of the most important factors in the success of the north sea basin as a major oil province is the presence of the thick succession of marine mudstones of late jurassic to earliest cretaceous age. this succession, which comes under various stratigraphic guises (figs 1, 2), is generally accepted to have been the source of much volgian–ryazanian ‘hot shales’ of the bo member (farsund formation) in the danish central graben, north sea: stratigraphy, facies and geochemistry jon r. ineson, jørgen a. bojesen-koefoed, karen dybkjær and lars h. nielsen upper jurassic – lowermost cretaceous marine mudstones represent the most significant source of hydrocarbons in the central and northern north sea. of particular importance in the danish sector of the central graben is a succession of radioactive ‘hot shales’ referred to the bo member, in the upper levels of the farsund formation (kimmeridge clay formation equivalent). this mudstone-dominated succession is typically 15–30 m thick and has a total organic carbon (toc) content of 3–8%, though locally exceeding 15%. although truncated on some structural highs, the bo member is a persistent feature of the danish central graben. lateral variation in both thickness and organic richness is attributed to intrabasinal structural topography and to the location of sediment input centres. detailed study of the dinoflagellate cyst biostratigraphy of 10 wells indicates that the onset of enhanced organic carbon burial began in the middle–late middle volgian in this portion of the central graben. the bo member, representing the peak of organic carbon enrichment, is largely of early ryazanian age. core data (jeppe-1, e-1 wells) indicate that the organic-rich shales of the bo member are not wholly of hemipelagic origin, as commonly assumed, but may locally be dominated by fine-grained turbidites. absence of bioturbation, well-preserved lamination and high toc values suggest that bottom waters were predominantly anoxic although the presence of in-situ benthic bivalves at discrete horizons in the e-1 well suggests that suboxic conditions prevailed on occasion. the bo member is a good to very good source rock, showing very high pyrolysis yields (10–100 kg hc/ton rock) and hydrogen index (hi) values in the range 200–600. in particular, the bo member is characterised by an abundance of 28,30 bisnorhopane (h28), a compound that is indicative of anoxic environments. these new data from the danish sector of the central graben are compatible with the model of tyson et al. (1979) in which the accumulation of organic-rich mudstones was controlled primarily by bottom-water anoxia beneath a stratified watermass. a number of factors probably contributed to the development of watermass stratification, both intrinsic such as the tectonic morphology of the graben system and extrinsic including climate and sea-level stand. keywords: danish central graben, north sea, volgian–ryazanian, organic-rich mudstones, marine source rock, depositional processes, geochemistry, anoxia geological survey of denmark and greenland, geocenter copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ji@geus.dk geological survey of denmark and greenland bulletin 1, 403–436 (2003) © geus, 2003 404 of the hydrocarbons in the central and northern north sea (barnard & cooper 1981; cornford 1994, 1998; kubala et al. 2003) and also forms the seal in many of the jurassic oilfields in this area (e.g. south brae, turner et al. 1987). in well sections, this basinal mud-dominated succession is characterised by alternations of ‘hot’ and ‘cold’ mudstone intervals (price et al. 1993). this refers to relative levels of radioactivity emitted by the mudstones, as recorded by the gamma-ray log. in general, such variation in the levels of gamma radioactivity in marine mudstones is assumed to reflect the organic richness of the sediment. of particular note in the central graben of the north sea are the ‘hot shales’ of volgian to ryazanian age which have been formalised as the clay deep member (kimmeridge clay formation) in the dutch sector and the mandal formation in the norwegian sector (fig. 2; vollset & doré 1984; van adrichem boogaert & kouwe 1993). in the danish sector, partially equivalent organic-rich mudstones have been described as the ‘hot unit’, an informal member of the farsund formation (jensen et al. 1986); this unit is formally defined as the bo member of the farsund formation in this volume (michelsen et al. 2003, this volume). in the uk sector of the central north sea, equivalent organic-rich mudstones are not given formal lithostratigraphic status within the kimmeridge clay formation (see the ‘hot shales’ of donovan et al. 1993). the aim of this paper is to integrate data from various ■■ ■■ ■ ■ ■■ ■■ ■ ■ ■ ■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■■ ■■ 0° 55°n 100 km outer moray firth v ik in g g ra be n sørvestlandet high central graben ringkøbing– fyn high n g nl uk dk danish basin basement or intrabasinal high english kimmeridge clay outcrop english kimmeridge clay subcrop international sector boundary normal fault■■ mid-north sea high east shetland platform fennoscandian shield 60°n 50°n 10°e fig. 1. generalised jurassic tectonic setting of the north sea region (based on doré et al. 1985). dk, denmark; g, germany; n, norway; nl, the netherlands; uk, united kingdom. disciplines and thus to provide an up-to-date assessment of the nature, distribution and origin of the ‘hot shales’ of the bo member in the danish central graben. geological setting the danish offshore area extends westwards to include a segment of the north sea central graben, a complex mesozoic rift system that trends roughly nnw–sse (fig. 1). although probably following palaeozoic lineaments (glennie 1990), a discrete rift system is thought to have first developed in the earliest triassic, trending north–south (ziegler 1988, 1990; sundsbø & megson 1993). the dominant nw–se structural trends in the danish central graben developed during the late jurassic when the rift was at its most active, with the accumulation of up to 4 km of (compacted) sediment in the most rapidly subsiding sub-basins of the danish central graben (møller 1986; sundsbø & megson 1993; japsen et al. 2003, this volume). the early cretaceous saw the transition from this phase of active extension and rapid fault-controlled subsidence in the late jurassic to the regional subsidence pattern, centred on the axial graben system, that characterised the late cretaceous and cenozoic. late jurassic – earliest cretaceous structural evolution the late jurassic was characterised by the development and pronounced differential subsidence of successive half-grabens within the complex rift basin of the central graben. subsidence and sedimentation was focussed on the eastern and southern area (søgne basin, tail end graben and salt dome province; fig. 3) during the latter part of the middle jurassic but extended north-westwards in the late jurassic due to both the overall sea-level rise and the development of secondary depocentres in successive half-grabens (møller 1986; damtoft et al. 1992; andsbjerg & dybkjær 2003, this volume; møller & rasmussen 2003, this volume). the feda graben in the danish sector was probably initiated in the early late jurassic (?oxfordian) and formed the dominant sediment depocentre in the area during the 405 northern dutch central graben danish central graben norwegian central graben uk central graben outer moray firth south viking graben (uk) age valanginian ryazanian volgian kimmeridgian oxfordian la te j ur as si c c re ta ce ou s vlieland claystone fm scruff greensand fm kimmeridge clay fm upper graben fm middle graben fm lola fm heno fm farsund fm bo member poul fm mandal fm ula fm farsund fm haugesund fm fulmar fm heather fm heather fm heather fm kimmeridge clay fm kimmeridge clay fm kimmeridge clay fm brae fm piper fm cromer knoll group cromer knoll group cromer knoll group cromer knoll group cromer knoll group clay deep mb fig. 2. generalised upper jurassic lithostratigraphy in the north sea rift system from the dutch sector of the central graben in the south to the moray firth and southern viking graben in the north (vollset & doré 1984; van adrichem boogaert & kouwe 1993; richards et al. 1993; lott & knox 1994; michelsen et al. 2003, this volume). sand-rich formations are indicated by a stipple ornament, lithostratigraphically defined ‘hot shales’ by oblique shading. 406 b a 56°00´n 4°00´e ? ? ? b-1 5 10 15 6 85 10 15 13 25 39 14 114 76 7 20 27 18 34 101 deep adda-1 elin-1 lone-1 kim-1 bo-1 edna-1 e-1 anne-3 56°00´n 4°00´e n. jens-1ravn-2 i-1 iris-1 v-1 ugle-1 20 km 20 km adda-1 jens-1 g-1 alma-1 s.e. igor-1 gwen-2 mid north sea high inge high c offee soil fault tail end g raben gulnare-1 salt d om e province gertrud graben feda graben eg-1 jeppe-1 deep gorm-1 gert-2 emma-1 ål basin gert ridge outer rough basin søgne basin mandal high ringkøbing–fyn high wells in which the bo member is present wells in which the bo member is not developed inferred distribution of the bo member marginal zone in which the bo member is probably not developed salt diapirs normal fault reverse fault kimmeridgian (söderström et al. 1991; johannessen & andsbjerg 1993; rasmussen 1995). the gertrud graben developed as a discrete depocentre during the late kimmeridgian whereas onlap onto the mid north sea high to the west began in the volgian (møller 1986; damtoft et al. 1992). in this western area, two discrete depocentres (the ål and outer rough basins) were initiated in the latest jurassic and formed important sediment sinks in the early cretaceous (japsen et al. 2003, this volume). during the latter part of the late jurassic, therefore, the danish central graben was segmented into a number of nw–se-trending depocentres, separated by elongate highs or broad plateaus (johannessen & andsbjerg 1993; andsbjerg & dybkjær 2003, this volume; johannessen 2003, this volume). this history of protracted extension during the middle and late jurassic was interrupted in the latest jurassic – earliest cretaceous by a complex tectonic phase that was essentially extensional in character but involved block rotation associated with localised compression, reverse faulting and uplift (rasmussen 1995; møller & rasmussen 2003, this volume). this end-jurassic tectonic phase coincided broadly with deposition of the organic-rich bo member and thus is of direct relevance to this study. stratigraphy history and stratigraphic status the organic-rich ‘hot shales’ of the bo member occur within the uppermost levels of the farsund formation (fig. 2). originally described informally by jensen et al. (1986) as the ‘hot unit’, the stratigraphy of this unit was subsequently discussed briefly by michelsen & wong (1991) and the hydrocarbon source rock characteristics were reported by østfeldt (1987) and damtoft et al. (1987, 1992). a detailed study of these deposits was undertaken by bojesen-koefoed (1988). as noted earlier, michelsen et al. (2003, this volume) formally define the ‘hot unit’ of jensen et al. (1986) as the bo member in an accompanying paper. log character the bo member is recognised primarily on the basis of log character since the gross lithological contrast between this member and the remainder of the farsund formation is often slight and rarely detectable in ditch cuttings. jensen et al. (1986) described the unit from the bo-1 well at the southern end of the tail end graben (figs 3, 4). this description relied heavily on the gamma-ray log which thus forms the essential criterion for its recognition. in bo-1, the background gamma-ray values of the farsund formation mudstones are in the range 75–100 api (fig. 4); the bo member shows values ranging from 120 to 160 api with marked upward shifts to higher and lower values defining the lower and upper boundaries of the member respectively. although generally high, gamma-ray values may vary significantly within the bo member; in bo-1 and many other wells, upward-decreasing gamma trends are evident, typically 3–5 m thick and separated by intervals showing more consistently high gamma-ray values. a marked feature of the bo-1 section is the upward increase in gamma-ray values beneath the bo member over an interval of about 50 m (8724–8561 ft). this ‘warming-upwards’ interval (henceforth referred to as the w-u interval) is also a feature of many other wells (see below). variation in the degree of development of this w-u interval in relation to the ‘hottest’ interval can result in ambiguity in locating the base of the bo member. in recognising the bo member throughout the danish central graben, the bo-1 gamma log pattern was used as the basic reference. it is important to note that the bo member of the danish sector of the central graben is more restricted in its definition than the partially equivalent mandal formation in the norwegian sector and the clay deep member of the dutch sector (fig. 2; dybkjær 1998). the bases of these units are defined at the point at which the overall gamma values begin to increase (i.e. at the base of 407 facing page: fig. 3. a: late jurassic tectonic framework of the danish central graben (modified from damtoft et al. 1992) showing the location of released wells in which the bo member is recognised. the bo member is not recognised in the wells indicated by open circles in the east of the area, although the same stratigraphic interval is represented in these wells. note that the ål and outer rough basins, although depicted on this map, were areas of active subsidence primarily in the early cretaceous (japsen et al. 2003, this volume). b: map showing the thickness (in metres) of the bo member in well sections and the lateral distribution of this member (blue) as deduced from well and seismic data. stipple indicates the inferred area in which the bo member is not developed due to siliciclastic dilution of the organic matter proximal to the ringkøbing–fyn high. the distribution of the bo member is not inferred for the southern area (salt dome province) due to sparse data points and the thin, irregular development of the ‘hot shales’ in this region. the w-u interval) and they are succeeded by cretaceous strata of the cromer knoll and rijnland groups (vollset & doré 1984; michelsen & wong 1991); the mandal formation and the clay deep member clearly span a greater stratigraphic interval than the bo member as defined by michelsen et al. (2003, this volume). spectral gamma-ray logs are not available for the bo-1 well but the example illustrated from the lone-1 well (fig. 5) shows an overall gamma log pattern that is comparable with the bo-1 section. it is clear from the spectral log that the increase in the overall gamma-ray values both beneath and within the bo member is the result of an increased content of uranium (jensen et al. 1986); the thorium and potassium values show little variation. as demonstrated by swanson (1961), uranium in sedimentary rocks is typically bound to organic matter and the positive correlation between gamma radioactivity and total organic carbon (toc) in the bo member is marked in most wells (see fig. 14; damtoft et al. 1992, fig. 5). uranium may also be concentrated in biogenic phosphate material, however, and since such debris is common in the organic-rich mudstones of the bo member, this may provide an additional contribution to the total gamma response. 408 gert-2 kim-1 gr (api) dt (µsec/ft) ild (ohms m2/m) dt (µsec/ft) ild (ohms m2/m) 132 0 150 13400 13600 13800 13200 14000 14200 13800 13600 0 8400 ft b.kb ft b.kb ft b.kb 8600 8800 bo-1 density (g/cm3) gr (api) gr (api) 1 10010150 120 600 150 ild (ohms m2/m) 10010 1 10 100120 60 fig. 4. the bo member (blue) in the bo-1 (the type well), gert-2 and kim-1 wells; for well locations, see fig. 3. the bo member is succeeded by the uppermost farsund formation in the bo-1 well, but is overlain directly by the cromer knoll group (åsgaard formation) in the gert-2 and kim-1 wells; the contact is probably a fault in the kim-1 well. the variation in log expression and stratigraphic development is discussed in the text. depth in feet below kelly bushing (kb); note that imperial units are retained for original well data measured in feet although metric units are preferred elsewhere. gr, gamma ray; dt, sonic velocity; ild, deep induction (resistivity). resistivity logs through this stratigraphic interval often show anomalously high readings which correspond, at least in part, to the bo member as defined on the gamma log. this gross correspondence is not surprising since resistivity logs are often used to recognise zones of organic richness (meyer & nederlof 1984; passey et al. 1990). in bo-1, the base of the bo member is marked by an abrupt increase in resistivity readings which remain relatively high throughout the unit (fig. 4). it is noteworthy, however, that these high resistivities continue upwards to the top of the farsund formation, above the upper boundary of the bo member as defined by the gamma log. in a number of wells, the resistivity logs display a stepped increase, corresponding roughly to the w-u interval on the gamma log beneath the bo member. it should be acknowledged, however, that the marked log shifts on the gamma and resistivity logs often do not coincide in detail, despite the broad correspondence, and the boundaries defined on the gamma log are considered to take priority, following jensen et al. (1986). in bo-1, the sonic and density–neutron logs do not vary substantially through this stratigraphic interval, indicating that gross lithological variation between the bo member and the host farsund formation is small; bulk densities and sonic velocities are slightly lower and neutron log values slightly higher in the ‘hottest’ zones of the bo member compared with the farsund formation in general. distribution and regional development the bo member is recognised widely in the danish central graben in wells where this portion of the upper farsund formation is preserved (fig. 3). the uppermost farsund formation is often truncated on structural highs (e.g. the inverted søgne basin, fig. 3) such that the original extent and variation in development of the bo member in the various subbasins is difficult to evaluate. furthermore, although forming part of the uppermost seismic sequence of the farsund formation mapped by møller (1986), the bo member alone cannot be differentiated on seismic data. knowledge of the lateral extent and variation in development of the member is thus fragmentary. it is possible, however, to make certain broad observations concerning both the distribution and lateral variation in the stratigraphic development of the bo member. firstly, it is absent in a number of wells flanking the eastern margin of the danish central graben, from gulnare-1 in the north to alma-1 in the south (fig. 3). in these wells, biostratigraphic data indicate the presence of the stratigraphic interval occupied by the bo member elsewhere yet the anomalously high gamma values characteristic of this member are not observed. it is likely that the lack of development of organic-rich mudstones in these marginal wells is largely the result of increased siliciclastic input and consequent dilution of the organic matter. it is noteworthy that discrete sand-rich intervals are characteristic of the uppermost jurassic – lowermost cretaceous in a number of these wells and indeed are recognised at formation 409 th k u 11600 11700 11800 11900 11500 ft.b.kb 0 150 -10 300 10 20 sgr (api) uranium (ppm) potassium (%) thorium (ppm) 0 0.05 0.1 0 20 40 c ro m er k no ll g p fa rs un d fo rm at io n bo m em be r fig. 5. spectral gamma log through the bo member in the lone-1 well. note that the increased gamma radioactivity exhibited by this interval is attributable solely to the contribution from uranium. sgr, total gamma-ray log; th, thorium; k, potassium; u, uranium. 410 kim-1 gert-2 jeppe-1 elin-1 no data gr 0 150 gr 0 150 gr 0 150 gr 0 180 6 5 2 6 5 4 1 2,3 4 5 4 3 1 1 5 6 1 nw core 100 m elin-1 kim-1 bo-1 edna-1 e-1 anne-3 20 km jeppe-1 gert-2 last occurrence datum, dinoflagellate cyst species: 6 dichadogonyaulax culmula/dingodinium? spinosum 5 rotosphaeropsis thula 4 amphorula expirata 3 egmontodinium polyplacophorum 2 dichadogonyaulax pannea/glossodinium dimorphum 1 senoniasphaera jurassica base cromer knoll group bo member base ryaz-1 sequence boundary base volg-4 sequence boundary fig. 6. variation in the development of the bo member in the danish central graben, illustrated by gamma logs (datum: top bo member). sequence boundaries ‘base volg-4’ and ‘base ryaz-1’ are from andsbjerg & dybkjær (2003, this volume). 411 level (poul and vyl formations; michelsen et al. 2003, this volume) in the deep adda-1, v-1 and ugle-1 wells. furthermore, the bo member is poorly defined, and only tentatively recognised, in several wells that lie adjacent to this marginal belt. in the elin-1 well in the tail end graben (fig. 3), for example, biostratigraphic data indicate that the upper volgian to ryazanian succession is highly expanded relative to other wells (fig. 6). the bo member in this well is thick (114 m compared with 39 m in bo-1) but shows only slightly higher gamma values than the background farsund formation. the thickness of the bo member varies greatly in the danish central graben, from less than 10 m in the southern salt dome province to over 100 m in the western part of the danish central graben (fig. 3). although this overall trend appears significant, thickness variation over much of the danish central graben does not show any systematic regional trend. it was probably controlled by local factors such as structural position within individual subbasins and local variations in the rate of sediment supply. the complexity is exemplified in the central portion of the danish central graben where the bo member is 39 m and 25 m thick, respectively, in the bo-1 and e-1 wells but under 15 m in nearby wells such as north jens-1 and jens-1 (fig. 3). the influence of structural position (and hence proximity to depocentres) is illustrated by comparing the gert-2 well on the gert ridge (18 m) with the jeppe-1 well within the gertrud graben (27 m) and the i-1 well in the major depocentre of the tail end graben (76 m). it should be emphasised again, however, that the bo member, although potentially thick, is only weakly developed in the major depocentre of the tail end graben. the log character of the bo member also varies through the danish central graben (figs 4, 6), often in association with variation in thickness. on structural bo-1 edna-1 e-1 anne-3 gr 0 180 gr 0 180 gr 0 180 gr 0 180 6 6 6 5 4 2 1 5 4 1 ? 5 4 2 1 5 4 3 2 1 se core 412 egmontodinium polyplacophorum dingodinium? spinosum daveya boresphaera, stiphosphaeridium arbustum amphorula expirata glossodinium dimorphum senoniasphaera jurassica scriniodinium inritibile, leptodinium subtile occisucysta balia rotosphaeropsis thula, systemaphora? daveyi albidum stenomphalus icenii kochi runctoni paratollia preplicomphalus lamplughi primitivus oppressus anguiformis kerberus okusensis glaucolithus albani fittoni rotunda boreal ammonite zonation u pp er ju ra ss ic lo w er va la ng . u pp er r ya za ni an lo w er r ya z. u pp er vo lg ia n up pe r m id dl e vo lg ia n m id dl e m id dl e vo lg ia n lo w er m id dl e vo lg ia n lo w er c re ta ce ou s chronostratigraphy lo w er m id dl e vo lg ia n va la ng in ia n to l. h au te ri vi an u pp er r ya za ni an lo w er r ya z. bo m em be r fa rs un d fm va lh al l fo rm at io n d.? spinosum daveya boresphaera s. arbustum r. thula s.? daveyi a. expirata g. dimorphum s. jurassica s. inritibile l. subtile o. balia e. polyplacophorum lo w er c re ta ce ou s u pp er ju ra ss ic fa rs un d fo rm at io n 8600 8200 8400 8800 0 50 100 150 chronostratigraphy lithostratigraphy depth (ft.b. kb) last occurrence datum (lod), dinocysts gamma-ray (api) a b last occurrence datum (lod), dinocysts u pp er vo lg ia n m id dl e m id dl e vo lg ia n up pe r m id dl e vo lg ia n fig. 7. a: biostratigraphy of the bo-1 well, the type well of the bo member, based on analysis of ditch cuttings. the stippled chronostratigraphic boundaries cannot be accurately positioned in the succession on the basis of the palynology (fig. 7b). b: correlation of dinocyst ‘tops’ to the standard boreal ammonite zonation compiled from costa & davey (1992) and riding & thomas (1992). highs, such as the gert ridge (gert-2, fig. 4), both the bo member and the underlying w-u interval may be thin in comparison to the type bo-1 well, although still well-defined on petrophysical logs. in a number of wells (e.g. jeppe-1, fig. 6), an important erosional surface has been recognised beneath the bo member, truncating the w-u interval and in places succeeded by deep-water sandstones. this surface has been interpreted as a sequence boundary of regional significance by andsbjerg & dybkjær (2003, this volume; see also below). biostratigraphy the farsund formation in the danish central graben has a maximum age range of kimmeridgian to ryazanian. the bo member (michelsen et al. 2003, this volume) occurs in the uppermost farsund formation, in the interval broadly dated as latest middle volgian to early ryazanian (poulsen 1991). previous work on both ammonites and dinoflagellate cysts (hereafter referred to as dinocysts) from the cored portion of the bo member in the e-1 well (fig. 6) indicated an early ryazanian age (kochi chronozone) for at least the uppermost bo member (birkelund et al. 1983). in the present study, the biostratigraphy of the bo member and the immediately underlying and overlying farsund formation was investigated by means of palynological analysis of core where available (e-1, jeppe-1), sidewall cores and ditch cuttings from a total of 10 wells. the results of this study are summarised in figures 6–8; detailed biostratigraphic data from bo-1 in comparison with the type well of the mandal formation (7/12-3a) have been presented by dybkjær (1998). a number of key dinocysts (fig. 8c–h) can be used to bracket the bo member over much of the danish central graben; these dinocysts are well-known as stratigraphically useful species in the north sea region (davey 1979, 1982; riding 1984; riding & thomas 1992). as illustrated for bo-1 (fig. 7), the top of the bo member typically occurs immediately above the last occurrence datum (lod) of rotosphaeropsis thula whereas the base falls between the lod of amphorula expirata and the lod of egmontodinium polyplacophorum. the interval showing the highest gamma values is bracketed by the lod’s of r. thula and a. expirata, corresponding to the kochi chronozone of the early ryazanian (fig. 7). the results from the bo-1 well, supported in general by the regional data (fig. 6), indicate a maximum age range for the bo member of late volgian (preplicomphalus chronozone) to middle late ryazanian (stenomphalus chronozone). the point at which the gamma values begin to increase, i.e. the base of the w-u interval underlying the bo member (equivalent to the base of the mandal formation in the norwegian sector), corresponds roughly to the lod of senoniasphaera jurassica in bo-1 and many other wells (figs 6, 7), giving a middle–late middle volgian age for the onset of increased preservation of organic carbon in the succession. over much of the danish central graben, the biostratigraphy of the bo member is closely comparable to that in bo-1 (fig. 6), despite the variation in thickness and local problems of definition (see discussion above). in most wells, the top of the bo member lies between the lod of dichadogonyaulax culmula/dingodinium? spinosum and that of r. thula as observed in the type well (figs 6, 7), although deviations from this pattern were recorded (compare gert-2 with e-1). over most of the transect in figure 6, from gert-2 in the north to anne-3 in the south, the base of the bo member coincides roughly with the lod of a. expirata. key volgian dinocysts are absent from the underlying farsund formation in the edna-1 and e-1 wells; in the latter case this may be attributable to an erosional hiatus near the base of the bo member (fig. 6; andsbjerg & dybkjær 2003, this volume). in the jeppe-1 well, the ranges of dichadogonyaulax? pannea/glossodinium dimorphum and e. polyplacophorum extend into the bo member and share a common lod (fig. 6). it is thought likely that this anomaly can be attributed to reworking of these species, extending their ranges upwards. as described in detail below, the strata underlying the bo member in jeppe-1 are characterised by abundant evidence of slumping and sediment gravity flow and the potential for redeposition of older sediments must be considered high. in the gert-2 well on the gert ridge, in contrast, these dinocysts are not recorded and uppermost middle to upper volgian strata may be absent in this well; a sequence boundary is recognised at this level in gert-2 by andsbjerg & dybkjær (2003, this volume; fig. 6). thus, the anomalous biostratigraphic results from this northern portion of the danish central graben are perhaps best explained by local erosion of structurally positive regions during the latest volgian with concomitant redeposition in adjacent lows; this subject is discussed further below. as noted earlier, the bo member is well-developed and anomalously thick in the kim-1 and b-1 wells in the westernmost part of the danish central graben (the incipient ål and outer rough basins). in addition, the 413 414 a c d e f g h b p d a a w w p p 100 µm 20 µm 20 µm 20 µm 50 µm 20 µm 20 µm 20 µm fig. 8. palynofacies assemblages (a, b) and stratigraphically important dinocysts (c–f) from the bo member and the upper farsund formation in general. the figured specimens are stored at the geological survey of denmark and greenland, copenhagen, under the catalogue numbers provided. a: a typical palynofacies assemblage from the organic-rich mudstones (facies 1, 2) of the bo member showing a dominance of aom (a), dinocysts (d) and prasinophyte algae (p). jeppe-1, core 1, 4402.71 m (drill depth). b: palynofacies assemblage from a muddy sandstone bed (facies 3b; fig. 13, sample 2). note the heterogeneous nature of the assemblage, comprising wood particles (w), aom (a) and palynomorph fragments, including prasinophyte algae (p). jeppe-1, core 1, 4418.73 m (drill depth). c: rotosphaeropsis thula. bo-1, cuttings sample 8670–8680 ft, geological survey of denmark and greenland (geus) catalogue no. 2000-kd-001. d: dingodinium? spinosum. bo-1, cuttings sample 8370–8380 ft, geus catalogue no. 2000-kd-002. e: systematophora? daveyi. bo-1, cuttings sample 8790–8800 ft, geus catalogue no. 2000-kd-003. f: amphorula expirata. bo-1, cuttings sample 8640–8650 ft, geus catalogue no. 2000-kd-004. g: egmontodinium polyplacophorum. bo-1, cuttings sample 8580– 8590 ft, geus catalogue no. 2000-kd-005. h: senoniasphaera jurassica. bo-1, cuttings sample 8760–8770 ft, geus catalogue no. 2000-kd-006. w-u interval that is characteristic of the strata underlying the bo member over much of the danish central graben is absent or very thin in this western area. the biostratigraphic data are poor but the results from kim-1 suggest that the lower portion of the bo member in this well may be time-equivalent to at least part of the w-u interval in the remainder of the danish central graben (fig. 6). sequence stratigraphic framework the sequence stratigraphy of the jurassic in the central graben presented by andsbjerg & dybkjær (2003, this volume) is adopted here. based primarily on new biostratigraphic data integrated with detailed log analysis, these workers subdivided the upper jurassic mudstonedominated farsund formation into 11 sequences. in their study, sequence boundaries are recognised in well sections on the basis of parasequence stacking patterns (derived from log analysis), abrupt facies shifts and biostratigraphic evidence for hiatuses. in most wells, the uppermost sequence boundary (base ryaz-1) lies beneath or at the base of the bo member (fig. 6). in wells exhibiting an expanded section (e.g. bo-1, fig. 6) this surface lies within the upper levels of the w-u interval and there is no evidence of a significant stratigraphic gap. in such apparently conformable sections, the sequence boundary is of late volgian or early ryazanian age. in other wells, particularly on late jurassic structural highs (e.g. the gert ridge), this sequence boundary lies close to, or is coincident with, the base of the bo member and is marked by a significant stratigraphic gap (e.g. gert-2, fig. 6). andsbjerg & dybkjær (2003, this volume) did not present a sequence stratigraphic interpretation of the succession overlying this sequence boundary up to the base of the cromer knoll group, acknowledging the probable complexity of the bo member itself and the unresolved debate concerning the sequence stratigraphic identity of the ‘base cretaceous unconformity’ (rawson & riley 1982; donovan et al. 1993). donovan et al. (1993) presented the results of an integrated sequence stratigraphic study of the middle–upper jurassic in a portion of the uk central graben, some 200 km north-west of the danish sector. they recognised eight depositional sequences, the uppermost of which includes volgian–ryazanian ‘hot shales’ that are equivalent in part to the bo member of the danish central graben. the lower boundary of this sequence (the ‘purple’ sequence boundary of donovan et al. 1993) occurs beneath the ‘hot shales’; in some wells, this is a major truncation surface at which the volgian and uppermost kimmeridgian are absent. at its correlative conformity, this surface is suggested to be of medial middle volgian age (base portlandian, sensu anglico). donovan et al. (1993) suggested further that the maximum gamma peak within the ‘hot shales’ represents a maximum flooding surface; the succeeding sequence boundary was placed at the base of the cromer knoll group although donovan et al. (1993) acknowledged that evidence for truncation at this surface is scarce. there is a striking similarity between the log character of the bo member in the danish sector and the equivalent succession in the uk sector of the central graben, particularly with respect to the relationship of these ‘hot shales’ to the regional sequence stratigraphic framework. however, the ages referred to the sequence boundaries underlying the ‘hot shales’ do not match and there is clearly a problem in integrating these two sequence stratigraphic frameworks. facies, processes and depositional environment the uppermost jurassic – lowermost cretaceous ‘hot shales’ in the north sea central graben occur within a thick upper jurassic – lower cretaceous mudstonedominated marine succession (farsund formation and cromer knoll group; fig. 2), an unattractive stratigraphic position with respect to exploration for hydrocarbon reservoirs. as a result, this interval is rarely cored (pegrum & spencer 1990) and lithological, biostratigraphic and geochemical studies are generally based on ditch cuttings and sidewall cores. the bo member of the danish central graben has, however, been partially cored in two wells, e-1 and jeppe-1 (figs 3, 6). the following discussion of facies, processes and depositional environment is based on these cores. it is wellestablished that the farsund formation was deposited in a fully marine environment (michelsen et al. 1987); the bo member yields ammonites, inoceramid bivalves and dinocysts confirming its overall marine character (birkelund et al. 1983 and this study). jeppe-1 well core 1 in jeppe-1 spans the lower boundary of the bo member (fig. 6), exhibiting 11.97 m of the underlying strata and 6.21 m of the bo member itself (figs 9–11). 415 416 c sl vf m sand f c vcc sl vf m sand f c vc c sl vf m sand f c vc 4407.5 4408.04405 depth in metres below kb 4410 4415 4416.0 4415.0 facies 1, 2: interbedded laminated mudstones and thin graded sandstone–mudstone couplets facies 3: structureless/graded sandstones facies 4: contorted sandstone–mudstone mudstone sandstone mudstone clasts contorted, slump-folded strata cross-lamination fig. 11a fig. 11b e d c b a 12 11 10 9 8 7 6 5 4 3 2 1 fig. 9. log of core 1 from the jeppe-1 well, spanning the base of the bo member (large arrow). note the clear subdivision into a heterogeneous, sandrich lower portion (up to base bo member) and an upper mud-dominated portion showing a fining-upwards trend. detailed logs illustrate the lithofacies described in the text; inferred sandstone injection structures occur at 4408 m and at 4415 m in these detailed sections. small arrows 1–12 indicate locations of palynofacies analyses exhibited in figure 13. the results of detailed lamina-bylamina logs (a–e) through the finegrained fraction are presented on table 1 and figure 10. the intervals illustrated in figure 11 are indicated. the core is well-preserved, in contrast to the e-1 core, and thus forms the dominant data source for detailed facies analysis, including palynofacies analysis. facies and depositional processes facies 1. black claystones this facies occurs intimately interbedded with thin-bedded sandstone–mudstone couplets (facies 2) where it ranges from laminae less than a millimetre thick to beds several centimetres thick. in representative detailed lamina-by-lamina sections (fig. 10), the black claystone laminae are typically 1–3 mm thick, increasing both in thickness and overall proportion of the section upwards from the sub-bo member strata into the bo member itself (fig. 12; table 1). laminae and beds are typically parallel-sided with sharp, flat boundaries, except where scoured and/or loaded at the contact with an overlying sandstone bed (facies 2). the claystones are black or very dark grey, locally with a faint brownish cast, and show a weak, yet pervasive planar structure defined by discontinuous organic wisps, an overall platy fabric and, in places, by concentrations of calcareous microfossils (largely rhaxella perforata, sponge reproductive cysts). bioturbation is absent. a characteristic feature of this facies is the occurrence of phosphatic fragments, typically concentrated along specific horizons; on bedding plane surfaces these are often identifiable as fish scales. larger vertebrate remains occur in this facies within the bo member at 4402.45 m (figs 9, 10); these have been preliminarily identified as vertebrae, ribs and other bones of a marine reptile, most likely of plesiosaur or ichthyosaur type (s.e. bendix-almgreen, personal communication 1994). due to the finely interstratified nature of facies 1 and 2, analysis of palynofacies and source rock potential was largely undertaken on mudstone plug samples that included both facies. as shown on figure 13 (see also fig. 8a), the composition of the organic matter in these composite samples is very uniform. amorphous organic matter (aom) and dinocysts together form over 75% of the organic matter, dominating over the terrestrial component (wood fragments, spores and pollen). the total organic carbon (toc) content of the mudstones (facies 1 and 2) in the bo member itself has a range of 5.2–7.1 wt% with an average of 6.0 wt% (table 2). in the succession beneath the bo member, the mudstones show toc values in the range 3.8–7.1 wt% (average of 5.0 wt%). supplementary analyses were undertaken on two samples of facies 1 and one sample of facies 2 from the bo member. these few analyses (table 2) suggest that the facies 1 mudstones show slightly higher toc values (7.3, 8.4 wt%) than the facies 2 mudstones (6.8 wt%), the slight difference being attributable largely to the higher silt content of the latter facies. interpretation. this facies is characterised by its finegrained nature, the high organic carbon content, the dominance of aom and dinocysts over terrestrial components, the lack of bioturbation and the weak yet pervasive planar fabric. all these features indicate deposition in a low-energy, oxygen-deficient marine environment, distant from terrestrial influence. the facies is interpreted to represent hemipelagic/pelagic fines, deposited by settling through the water column. the preservation of very thin (millimetric) sedimentation units in this succession, as a result of the suboxic–anoxic bottom conditions (terminology from tyson & pearson 1991) and consequent absence of bioturbation, permits the identification and differentiation of hemipelagic deposits from fine-grained sediment gravity flow deposits (see facies 2); in many deep-water fine-grained deposits, such differentiation is impossible (e.g. ineson 1989). this facies is equivalent to the ‘fissile-laminated non-bioturbated mudrock’ facies of stow & atkin (1987), in their study of upper jurassic mudrocks from the uk sector of the north sea, and to facies e2.2 of the deepwater facies scheme of pickering et al. (1986). facies 2. sandstone–mudstone couplets this facies is an important component of the sub-bo member succession and forms over 70% of the cored portion of the bo member itself (table 1). where fully developed (figs 9–11), it comprises two well-differentiated components: a lower parallelto cross-laminated, fineto very fine-grained sandstone succeeded by a structureless or coarse-tail graded, weakly laminated silty mudstone. such couplets may be up to 5 cm thick, but are typically 0.5–2 cm thick; the sandstone portion is typically in the range 1–5 mm whereas the mudstone portion is typically 5–10 mm thick. this facies, together with facies 1, is a common component of the slump sheets assigned to facies 4. syndepositional extensional microfaults and boudinage are observed locally. the organic composition of facies 2 is discussed under facies 1 (see above). the basal sand–silt layer has a sharp base, often erosional and loaded. some of these sandstones are normally graded but most are ungraded and show parallel417 418 section d section b 2 cm hemipelagic mudstone claystone cap structureless silty mudstone, weak coarse-tail grading laminated silty/ sandy mudstone cross-laminated sandstone fig. 10. logs of the fine-grained fraction (facies 1, 2) from the lower levels of the bo member (section b) and from the upper part of the cored portion of the bo member (section d) in jeppe-1 (for precise locations, see fig. 9). core photograph shows the thinly-bedded mud-dominated nature of the bo member (section d); arrows indicate the basal contacts of four typical sandstone–mudstone couplets (facies 2), interbedded with thin (1–3 mm) laminae of hemipelagic mudstone (facies 1). note the vertebrate remains (?plesiosaur/ichthyosaur) in the upper part of section d. inset shows an idealised fine-grained turbidite in the jeppe-1 core. or cross-lamination. such cross-lamination is low-angle, sometimes with muddy toesets, and the sandstones are often lenticular, resembling the ‘fading ripples’ of stow & shanmugam (1980). such well-developed sand–mud couplets (sand component > 2 mm thick) are subordinate in the facies (under 20% of facies 2 beds in the detailed sections) and many beds (about 30% in the detailed sections; fig. 10; table 1) possess only a discontinuous sand or silt lamina, less than 2 mm thick. in the bo member itself, nearly 70% of facies 2 beds lack the sand portion (i.e. ‘base cut-off’). the mudstone portion of these couplets typically abruptly overlies the sand component, if present, although graded transitions were observed in a few beds. the silty mudstones are commonly structureless but may show weakly defined coarse-tail grading, particularly in the lower levels in association with diffuse lamination. on polished slabs, the grading is picked out by an upward decrease in the proportion of dispersed very fine sand and coarse silt grains, or rhaxella cysts at some levels in the core. an interesting feature of this facies is the presence of a very thin (0.1–0.2 mm) but persistent cap of midbrown claystone (fig. 10). in detailed lamina-by-lamina sections, over 75% of beds with non-erosional upper contacts possess this ultra-thin claystone cap. the base of this layer is sharp although rapid grading from dark grey silty mudstone to paler-coloured claystone is observed; the upper contact with succeeding hemipelagic deposits (facies 1) is sharp and planar. interpretation. although the two components of these sandstone–mudstone couplets are often well-differentiated, their close association and the occurrence, in places, of a graded transition from sand to mud indicates that they are the result of a common process. the scoured basal contacts and the presence of grading indicates deposition from a waning turbulent bottom current. in the absence of evidence of wave or storm activity, and in view of the close similarity to descrip419 facies 1 hemipelagic mud facies 2 sand–mud couplets (fine-grained turbidites) e 27.3 2.4 (n = 33) 28.9 3.8 (n = 50) 60.4 10.7 d 26.5 2.7 (n = 31) 31.9 4.6 (n = 39) 64.6 3.5 c 25.4 3.1 (n = 16) 16.8 7.9 (n = 26) 76.5 6.7 b 35.9 1.5 (n = 18) 7.6 12.3 (n = 27) 73.4 19.0 a 53.0 1.1 (n = 12) 2.4 16.2 (n = 32) 81.9 15.7 proportion of section (%) average thickness (mm) average thickness (mm) proportion of mud component in section (%) proportion of sand component in section (%) section thickness (cm) table 1. jeppe-1 well, core 1: sedimentological data from the bo member (sections b–e), and underlying strata (section a) depth (m) facies toc (wt%) tmax (°c) s1 (mg/g) s2 (mg/g) hi 4401.36 1, 2 5.5 445 4.5 28.0 508 4401.37 1 7.3 446 3.6 33.0 454 4402.67 1 8.4 445 4.5 39.5 473 4402.69 2 6.8 444 4.0 31.2 458 4402.71 1, 2 7.1 444 5.2 37.7 533 4404.18 1, 2 5.2 443 4.1 24.7 477 4404.97 1, 2 6.6 441 5.5 35.7 543 12 4406.98 1, 2 5.9 443 3.6 28.7 488 11 4407.19 1, 2 5.6 444 3.0 27.9 497 10 4408.09 1, 2 4.7 441 2.5 20.8 443 9 4409.35 1, 2 3.8 441 2.6 14.5 378 4410.86 1, 2 5.3 443 2.6 22.4 425 4411.12 3a 0.5 434 2.0 0.8 169 8 4411.36 3a n.d. n.d. n.d. n.d. n.d. 7 4412.94 1, 2 7.1 433 5.4 36.9 518 4414.72 1, 2 5.3 445 3.1 22.6 431 4414.97 3b 0.2 n.d. n.d. n.d. n.d. 6 4415.49 1, 2 4.7 445 2.6 19.3 412 5 4417.33 1, 2 4.5 442 2.9 18.3 405 4 4417.47 4 2.3 443 2.3 6.3 276 3 4418.73 3b 1.2 442 1.7 3.5 281 2 4419.36 4 4.7 445 2.7 22.0 466 1 n.d., not detected palynofacies sample no. table 2. jeppe-1 well: geochemical data from core 1 spanning the base of the bo member (dashed line) 420 tions of fine-grained sediment gravity flow deposits in the literature (piper 1972; stow 1979; stow & shanmugan 1980; pickering et al. 1986), these beds are attributed to deposition from sediment-starved, dilute turbidity currents. the basal laminated and ripple cross-laminated sandstone portion, where present, is indicative of tractional processes whereas the succeeding structureless or graded mudstone portion was deposited prifig. 11. selected intervals of core 1, jeppe-1 (fig. 9); the section youngs from bottom left to top right. depths (drill) in metres below well reference point. a: the lowermost beds (4419.4–4416.9 m) typify the cored interval beneath the bo member. massive sandstones (s; facies 3) and slump sheets (ss; facies 4) are interbedded with fine-grained turbidites and hemipelagic mudstones (facies 1, 2). b: this core section (4408.9–4406.1 m) spans the lower boundary of the bo member (arrow; see also fig. 9) and illustrates the thinbedded, mud-dominated nature of this member. s a b ss 20 cm 0 421 marily from suspension. the silty lamination observed locally within the mudstones may be the result of shear sorting within the bottom boundary layer (stow & bowen 1978). these couplets are comparable to facies c2.3/d2.1 of pickering et al. (1986) and closely resemble the ‘mudstone facies’ described from the upper jurassic brae oilfield of the uk sector of the north sea (stow et al. 1982). these beds can be described as tc(d)e turbidites on the classical bouma (1962) scheme; within the mudstone portion, the divisions e1, e2, e3 of piper (1972) and t3/t4, t6, t7 of stow & shanmugan (1980) are represented. interestingly, the discrete, ultra-thin (0.1–0.2 mm) claystone cap observed in this study is not evident in these facies schemes for fine-grained turbidites. this cap is clearly differentiated from the graded or structureless mudstone division and from overlying hemipelagic deposits; it probably represents the fall-out of the finest sediment fraction following the passage of the muddy turbulent cloud. preservation of this thin lamina requires the total absence of burrowing infauna, perhaps explaining its lack of recognition in other finegrained successions. as noted under facies 1, the absence of bioturbation in association with the high toc content and the nature of the palynofacies indicates deposition in an oxygen-deficient marine environment. facies 3. sandstones this facies is absent from the cored portion of the bo member but forms about 17% of the underlying farsund formation (fig. 9). the sandstones are typically mediumto fine-grained, although locally coarseor very coarsegrained, and commonly contain up to 20% abraded shell fragments (bivalves, echinoderms, sponge spicules, bryozoans). the sandstones are mineralogically submature, containing up to 30% feldspar grains, mica flakes and rock fragments. the lithics include basement lithologies, such as meta-quartzite, acid plutonics and foliated mica-quartz aggregates, in addition to volcanic and dolomite fragments probably derived from the permian section. coalified wood fragments, locally up to 1 cm across, are present in some beds. glaucony is a characteristic although minor component. the toc content of this facies varies considerably, dependent on the proportion of mud either as matrix or as discrete clasts (table 2). two subfacies are recognised, based on the presence or absence of mud matrix. subfacies 3a. these mud-free, calcite-cemented sandstones are rare, being represented by just three beds ranging in thickness from 5 cm to 20 cm. they are of medium or medium–coarse sand grade and are structureless or show diffuse parallel to low-angle (deformed?) lamination. bed boundaries are typically sharp and planar but one bed has an erosional, scoured base and a normally graded, pebbly basal few centimetres (fig. 9). this facies occurs in association with muddy sandstones (facies 3b) and slumped sediment (facies 4). interpretation. the sharp, locally erosive basal contacts, the presence of normal grading, the massive to parallel-stratified structure and the lack of mud matrix suggests deposition from energetic, waning currents that were capable of sorting sand from mud-grade sediment. these beds are interpreted as the deposits of sandy turbidity currents and can be broadly classified as bouma tab turbidites. the implications of their close association with facies 3b and 4 is discussed below. the 40 15 10 5 0 30 20 10 0 e d c b a e d c b a average turbidite thickness (mm) proportion of hemipelagic mud (%) d et ai le d s ec tio ns d et ai le d s ec tio ns fig. 12. histograms showing the stratigraphic variation in the average thickness of the fine-grained turbidites (facies 2) and in the relative importance of the hemipelagic deposits (facies 1), based on the detailed sections a–e (fig. 9). 422 organic content of this subfacies shows a typically marine signature, aom and dinocysts together dominating the assemblage (fig. 13). subfacies 3b. mud-rich, mediumto coarse-grained sandstones dominate facies 3 and occur in close association with slumped sediments (facies 4). indeed, there is a complete gradation between muddy sandstones containing discrete intraformational sandstone or mudstone clasts and intervals of contorted sandstone and mudstone (cf. in situ facies 1/2) in which semi-coherent, slump-folded sediment rafts are separated by zones of muddy sandstone (cf. facies 3b). sandstone beds referred to this subfacies are 5–25 cm in thickness (typically about 10 cm thick) and generally show sharp, flat bed boundaries; loading and water escape structures are observed in places. they are typically structureless but a few beds show faint parallel stratification. elongate intrabasinal clasts ranging in length from a few centimetres to the width of the core (10 cm) are oriented parallel to bedding and in several cases are concentrated in the upper levels of the bed. two muddy sandstone beds assigned to this facies were subjected to palynofacies analysis (fig. 13, samples 2, 6); the results contrast with the fine-grained fraction (see discussion under facies 1) but also differ from each other. both samples show a decrease in the relative importance of aom compared with the mudstone 12 11 10 9 8 7 6 5 4 3 2 1 1, 2 1, 2 1, 2 1, 2 3a 3a 3b 1, 2 1, 2 4 3b 4 100% aom (amorphous organic matter) wood dinoflagellate cysts prasinophyte algae spores and pollen palynofacies category faciessample no. fig. 13. selected results of the palynofacies study of the jeppe-1 core (for sample locations, see fig. 9); a total of 20 samples were subjected to palynofacies analysis and a minimum of 500 kerogen particles were counted in each sample. note the consistent results obtained from the fine-grained fraction (facies 1, 2), and the atypical composition of the organic matter from the muddy sandstones (facies 3b). dashed line indicates the base of the bo member; note that the organic composition of the bo member mudstones is closely comparable to that of the mudstones in the underlying strata. facies, but the lower sample (fig. 8b) shows a relative increase in the wood component whereas the upper sample shows an increase in the proportion of dinocysts and prasinophyte algae. interpretation. the poor sorting, mud matrix, structureless ungraded character, flat non-erosional boundaries and gradation to slumped strata indicate deposition from viscous sediment gravity flows of debris flow type. this facies is equivalent to facies c1.1 of pickering et al. (1986). a few thin beds (max. 10 cm thick) assigned broadly to this facies show some features that are not wholly compatible with the interpretation given above (fig. 9). they show sharp, subparallel boundaries that are locally slightly oblique to the general bedding and display elongate, tapering offshoots up to a few centimetres long and a centimetre across, both on the upper and lower bed boundaries. although largely structureless, weak lamination may be present in the middle zone of the bed. these sandstones are interpreted to be of intrusive origin (i.e. sandstone sills) rather than representing primary sediment gravity flow deposits, although confirmation is impossible in core. facies 4. contorted sandstone–mudstone units of contorted intraformational sediment form an important and striking part of the sub-bo member succession in the jeppe-1 core (figs 9, 11). they are 10–50 cm thick and typically show flat, non-erosional bases and flat or slightly irregular tops. a complete gradation is represented from sheets composed entirely of slumpfolded but essentially coherent thinly interbedded sandstones and mudstones (cf. facies 1, 2) to sheets composed of lenses (phacoids sensu voigt 1962) of internally deformed bedded sandstone and mudstone floating in a muddy sandstone matrix. with increasing disintegration of intraformational slabs, this facies grades into the muddy sandstones assigned to facies 3b. the organic composition and toc content of this facies are very variable, most likely due to the variable proportion of mud in these heterogeneous deposits. interpretation. this facies records remobilisation and partial disaggregation of thin-bedded sandstone and mudstone and is thought to have originated largely by surficial downslope transport by processes ranging between sliding or slumping and viscous debris flow. it is thus equivalent to facies f2.1 transitional to c1.1 of pickering et al. 1986. it is possible, however, that subsurface injection processes may have created some of the fabrics illustrated by this facies (cf. anderton 1997). stratigraphic distribution of facies as shown on figure 9, the cored section in the jeppe-1 well is readily subdivided into a lower heterogeneous, relatively sand-rich interval capped by a fining-upwards, mud-dominated unit, representing the basal beds of the bo member. the base of the sand-rich interval beneath the bo member occurs at a log depth of 4431 m, at an inferred sequence boundary (fig. 6; andsbjerg & dybkjær 2003, this volume). in the cored sub-bo member interval, packets of thinbedded sandstone–mudstone turbidites (facies 2) with intervening hemipelagic laminae and beds (facies 1) make up around half (47%) of the succession, interbedded with sandstone turbidites and debris flows (facies 3; 17%) and sandstone–mudstone slump sheets (36%). in a representative detailed section (section a) through the fine-grained fraction (figs 9, 12; table 1), hemipelagic mud (facies 1) forms less than 5% of the fine-grained fraction at this level, typically occurring as thin laminae (about 1 mm thick) sandwiched between sand–mud turbidites (facies 2; average thickness 16 mm). the cored section of the bo member is composed solely of facies 1 and 2. the abundance and thickness of the sand component in facies 2 decreases upwards from the sub-bo member succession into the bo member itself, in parallel with an upward decrease in turbidite thickness and a relative increase in the proportion of hemipelagic mud (figs 10, 12; table 1). in the uppermost detailed sections (fig. 9, sections d, e), hemipelagic mudstone forms about 30% of the succession in laminae up to 10 mm thick. depositional setting the cored section from the jeppe-1 well records deposition in a low-energy marine environment characterised by background sedimentation of muds and subordinate thin sands from dilute, muddy turbidity currents and by suspension settling through the water column. the absence of bioturbation and the preservation of high levels of organic carbon indicate very low levels of free oxygen (suboxic–anoxic) within the bottom waters. in contrast, coarse poorly-sorted shelly sandstone turbidites, debris flow deposits and slump sheets record a more dynamic depositional environment. firstly, the 423 shelly sands testify to a source of immature sediment, including basement rock fragments; the presence of a varied assemblage of coarse shell fragments and glaucony suggest relatively high-energy, shallow marine conditions in the source area. secondly, the slump sheets of turbiditic sandstone and mudstone and associated debris flow deposits provide evidence of intrabasinal slopes in the vicinity of the jeppe-1 well. indeed, the sandstone injection structures, the evidence of minor slope creep and the close similarity between the in situ sediments (facies 1, 2) and the slump sheet components indicate that this succession accumulated very close to the base-of-slope. the jeppe-1 well is located near the western margin of the gertrud graben, a half-graben that was most active in volgian times when it was bounded by the gert ridge to the west and the mandal high to the east (fig. 3). as noted earlier, block rotation in association with subsidence of the feda and gertrud grabens resulted in local compression and uplift of the gert ridge during the latest jurassic (rasmussen 1995), yielding a potential sediment source just west of the jeppe-1 location. in the gert field at the northern end of the gert ridge, permian, carboniferous and metamorphic basement were encountered beneath the jurassic succession (rasmussen 1995). in the immediate area of the field, such potential sediment sources were draped by upper jurassic sediment during deposition of the bo member, but it is possible that these strata were locally exposed in uplifted fault slices at the southern end of the gert ridge. it is likely, however, that the gert ridge was most important as a source of intraformational sediment in the latest jurassic and earliest cretaceous. both the sedimentological and palynological data from jeppe-1 testify to significant reworking of middle to upper volgian sediments in the form of slumps, slides and debris flows. the well data from the gert field atop the gert ridge also suggest that the ridge experienced significant erosion in the latest jurassic. the uppermost farsund formation, including the bo member, is absent from the gert-1 and gert-3 wells and a significant stratigraphic gap spanning the late middle and late volgian is recognised immediately beneath the bo member in gert-2 (fig. 6 and previous discussion). the data suggest, therefore, that the cored succession in jeppe-1 was derived from two sources. the thin muddy turbidites that form the bulk of the succession were probably derived from regional sediment sources; transport paths are difficult to infer, especially given the complex nature of the danish central graben in the late jurassic, but were most likely axial in the various elongate subbasins. in this context, it should be noted that rasmussen et al. (1999) have proposed, on the basis of seismic data, the existence of a channelised sandy fan system in the axis of the gertrud graben that was broadly coeval with the cored section beneath the bo member in jeppe-1. these workers suggested that the thin-bedded turbidites observed in the jeppe-1 core represent ‘fan-fringe’ or levee/overbank deposits related to the axial channel system. rasmussen et al. (1999) postulated that the siliciclastic source for this fan system was the sørvestlandet high (fig. 1), although the mandal high and the inverted søgne basin area are also possible candidates (fig. 3). it is notable, however, that the mudstone component of the thin turbidites is palynologically and geochemically closely comparable to the hemipelagic mudstones. this suggests that the clay fraction was mainly of intrabasinal origin, cannibalised by erosive, turbulent flows entering the central graben. in contrast, the coarse shelly sands and the slumps, slides and debris flow deposits of intraformational sediment may have been of local origin, perhaps shed from the gert ridge immediately to the east of the gertrud graben. dispersal of this intrabasinal sediment from the flanks of the gert ridge may have occurred during storms; the close association of shelly sands with slump sheets or debris flow deposits suggests a common triggering mechanism. although acknowledging the possibility of local derivation, rasmussen et al. (1999) suggested that the shelly sands and slump–debris flow deposits may alternatively be related to the axial fan system, perhaps recording periodic levee collapse or breach. a number of the facies displayed by the jeppe-1 well are closely comparable to those described from the brae and miller oilfields in the viking graben of the north sea (stow et al. 1982; turner et al. 1987; mcclure & brown 1992). in particular, the sandstone–mudstone couplets (facies 2) are closely comparable to the ‘tiger stripe’ facies described by stow et al. (1982). such facies are commonly termed ‘interchannel’ or ‘levee’ deposits when observed in close association with coarse-grained channelised turbidites (mutti 1977; walker 1985). the inferred existence of a channelised fan system in the axis of the gertrud graben (rasmussen et al. 1999) is interesting in this respect. the large-scale fining-upwards trends observed in the core and in the overlying uncored portion of the bo member (fig. 6), record pulses of erosion and sediment dispersal both from regional and intrabasinal sediment sources. as discussed above, the mud-rich turbidites in the cored section become thinner, and in general finer-grained, upwards whereas the 424 interbedded hemipelagic muds form an increasing proportion of the succession (fig. 12; table 1). although the biostratigraphic resolution does not permit direct measurement of sedimentation rates at this scale, this pattern is suggestive of a waning supply of turbiditic mud relative to the background hemipelagic rain, perhaps related to rising sea level. e-1 well the cored section in the e-1 well is from the uppermost portion of the bo member, where the gamma-ray values are consistently high (fig. 6). although originally totalling some 2 m of core (9783–9792 ft, 77% recovery), this core dates from 1968 and has been intensively sampled. representative slabs remain, totalling about 1 m of core. this core and the boundary between the farsund formation and the overlying cromer knoll group in this well were subjected to a detailed biostratigraphic study by birkelund et al. (1983). the core fragments that remain are composed solely of parallellaminated black or very dark grey claystone, comparable to facies 1 of the jeppe-1 core. facies and depositional processes facies 1. black claystones the lamination in this dark organic-rich claystone is defined by slight colour variation, concentrations of silt-sized calcite grains (?rhaxella sp.) and phosphatic fish fragments. birkelund et al. (1983; fig. 3) illustrated the well-laminated nature of this facies by means of an x-radiograph, and suggested that two orders of lamination are present. the sub-millimetric parallel lamination is interrupted at regular intervals (5–20 mm) by discrete paler homogeneous claystone laminae (0.1–0.3 mm thick); these laminae define the second order lamination of birkelund et al. (1983). ammonites and inoceramid bivalves are present in this core; the latter occur at discrete horizons and in places appear to extend across the full width of the core. birkelund et al. (1983) figured complete inoceramid valves (assigned to inoceramus aff. vereshagini) and cross-sections indicate the presence of compacted yet entire valves; no evidence of transport is observed. the organic composition of the bo member in the e-1 core is very uniform (table 3); toc values range from 6.7 to 8.5 wt% (average of 7.3 wt%), comparable to the individual analyses of the hemipelagic facies (facies 1) in jeppe-1. the content of organic matter is also very homogeneous, being strongly dominated by amorphous organic matter (aom), dinocysts and prasinophyte algae. interpretation. in common with facies 1 of the jeppe-1 core, this well-laminated organic-rich mudstone facies is attributed to hemipelagic settling of fines through the water column. the discrete light-coloured claystone laminae are reminiscent of those that cap the mudstone turbidites (facies 2) in the jeppe-1 core and may have a similar origin i.e. they may represent the fine-grained tail of individual turbidite events. sandstone–mudstone couplets of the type seen in jeppe-1 (facies 2) are, however, not recognised in e-1. the well-preserved parallel lamination with no sign of bioturbation (even in radiographs), and the high content of organic matter (toc), particularly aom, suggest that anoxic conditions prevailed for much of the time in the bottom waters. the occurrence of in-situ inoceramid bivalves at certain horizons, however, indicates at least periodic suboxic conditions on the sea floor (birkelund et al. 1983). bioturbation is not evident in association with these faunas suggesting that conditions on the sea floor were close to the boundary between anoxic and dysoxic (i.e. suboxic in the terminology of tyson & pearson 1991). comparable apparently anomalous sediment–faunal associations have been described and similarly interpreted by savrda & bottjer (1987) from the miocene of california and by doyle & whitham (1991) from the upper jurassic – lower cretaceous of the antarctic peninsula. savrda & bottjer (1987) suggested that the development of such associations (their exaerobic biofacies) is favoured by the development of bacterial mats on the sediment surface (see also tyson & pearson 1991). the association of benthic inoceramid faunas, in particular, with laminated ‘black shales’ has been noted by many workers (see discussion by macleod & hoppe 1992). on the basis of facies criteria (kauffman & 425 depth (ft) facies toc (wt%) tmax (°c) s1 (mg/g) s2 (mg/g) hi 9784 1 7.0 431 2.7 40.6 582 9786 1 6.8 424 4.0 40.6 601 9787 1 7.0 426 3.5 42.9 614 9788 1 7.8 426 4.4 45.5 587 9789 1 6.7 433 3.6 40.2 599 9790 1 8.5 430 4.8 56.3 660 table 3. e-1 well: geochemical data from the bo member in core 8 sageman 1990) and isotope data (macleod & hoppe 1992), it has been suggested that inoceramids may have benefited from bacterial chemosymbiosis, thus extending their potential environmental range. although the isotopic evidence has been disputed (grossman 1993), there is well-documented evidence that inoceramid bivalves were tolerant of conditions that excluded most other forms (macleod & hoppe 1992), whether this was the result of a highly efficient metabolism, chemosymbiosis or a combination of these factors. in any event, inoceramids are not observed in the jeppe-1 core and only occur at specific levels in the e-1 core, indicating that even the most tolerant benthic invertebrates were largely excluded from the floor of the danish central graben during the deposition of the bo member. depositional setting the e-1 core consists solely of laminated organic-rich hemipelagic mudstones that accumulated under suboxic–anoxic bottom conditions. in contrast to the jeppe-1 core, turbidite processes were apparently unimportant in this setting. as shown on figure 6, however, the e-1 and jeppe-1 cored sections are not time-equivalent so that it is not clear whether this contrast in sedimentation style reflects a geographical or a temporal shift in the dominant depositional processes. organic geochemistry and source rock potential a number of studies have documented the good to very good source potential of the farsund formation in the danish sector (damtoft et al. 1987, 1992) and a positive correlation between source characteristics and the produced hydrocarbons in danish fields has been achieved in a number of cases (østfeldt 1987; unpublished geus data). in this section, the source characteristics and geochemistry of the farsund formation are described with particular emphasis on the bo member. farsund formation the total organic content (toc) of the farsund formation is very variable, being dependent on lithology, stratigraphic position, geographical setting and level of thermal maturity. it ranges from less than 1 wt%, typically in sandstone or dolomite/limestone stringers and in mudstones in the lower levels of the formation, to more than 15 wt% in the mudstones of the bo member. pyrolysis yields (rock-eval s2) vary from less than 1 kg hc/ton rock to more than 90 kg hc/ton rock. corresponding values of the hydrogen index range from less than 100 to approximately 600. however, in general, the farsund formation can be considered a good or even very good hydrocarbon source rock. the source rock potential varies with depth through the formation, as well as geographically within the central graben. in general terms, the lower farsund formation is poorer in organic carbon and the proportion of terrigenous organic matter is significant, leading to a mixed gas-/oil-prone kerogen type (damtoft et al. 1987, 1992). upwards, the terrigenous component of the organic matter decreases in abundance, leading to highly oil-prone kerogen of predominantly marine/bacterial origin. in the south-eastern part of the central graben, and along the eastern border fault (coffee soil fault), the source rock potential deteriorates, probably due to dilution with mineral matter (siliciclastic detritus) and incorporation of larger proportions of inert terrigenous organic matter. in the upper farsund formation, n-alkane distributions are unimodal, with low to moderate proportions of ‘unresolved complex mixture’ (ucm), centred in the range c15–19. with a few exceptions, the n-alkane distributions are smooth, with little or no preference for odd or even numbered compounds. the abundance of linear isoprenoids is generally low to moderate. pristane/phytane ratios are mainly in the range 0.9–1.6. occasionally, an increased contribution of terrigenous organic matter to the kerogen is manifest in slightly increased abundance of waxy components (nc22+), slight predominance of odd-numbered n-alkanes in the range nc23–31, and a bimodal distribution of ucm. in the lower farsund formation these features are pronounced, testifying to a general increase in the proportion of the kerogen component derived from terrigenous organic matter. the terpane distributions show very variable, but mostly modest amounts of tricyclic triterpanes, which may form a homologous series ranging from c20 to c30 (see below). the pentacyclic terpanes of the hopane series are dominated by hopane and norhopane. in low maturity samples, moretanes and ßß-hopanes, in particular 17ß(h)-trisnorhopane, may be rather abundant, whereas 28,30-bisnorhopane, where present, forms only a minor proportion. the proportions of ts and tm, 426 427 and 29ts vary with the level of thermal maturity. c30 diahopane (compound ‘x’ of philp & gilbert 1986; moldowan et al. 1991), is generally present in very low proportions, but tends to become increasingly prominent with maturation. extended hopanes are abundant, displaying a regular decrease in abundance with increasing carbon number from c31 to c35. the distribution of regular steranes is very homogenous, featuring a slight predominance of c27 steranes over c28 and c29 steranes, which are roughly equal in abundance. diasteranes are rather abundant, generally increasing with level of thermal maturity. c30 steranes are present in all samples. bo member at ‘bulk level’, the mudstones of the bo member are characterised by high or even very high organic carbon contents, generally in the range 4–8 wt% toc, occasionally exceeding 15 wt% toc. pyrolysis yields are very high, 10–100 kg hc/ton rock (rock-eval s2), with corresponding values of the hydrogen index occasionally exceeding 500, somewhat dependent on the level of thermal maturity (figs 14, 15; tables 2, 3). interestingly, organic matter enrichment may extend above the upper boundary of the bo member as defined by the gamma log (ravn-2, g-1). as discussed earlier, the high gamma radiation exhibited by the bo member mudstones is attributed primarily to uranium bound to organic matter. the occurrence of organic-rich shales with relatively low gamma values above the bo member may, therefore, be either the result of exhaustion of available uranium in the geochemical system or a decrease in the ability of the sediments to incorporate uranium. factors that may conceivably influence the incorporation of uranium into sediments are organic matter type, redox conditions and sedimentation rate. the data show no marked changes in organic matter type and, in any event, complexation and subsequent reduction of u6+ 10000 10050 10100 10150 10200 10250 10300 10350 10400 0 50 100 150 0.0 2.0 4.0 6.0 8.0 0 200 400 600 0.1 0.2 0.3 0.4 0.0 0.5 1.0 1.5 0.0 0.3 0.6 0.9 1.2 0.75 1.00 1.25 1.50 gamma ray (api) toc (%) hi t23/h30 h28/h29 h35/h34 cpidepth (ft.b.kb) fig. 14. geochemical profile of the upper farsund formation, including the bo member (blue), in the edna-1 well. toc, total organic carbon; hi, hydrogen index; t23/h30, ratio of c23-tricyclic terpane to c30 hopane; h28/h29, ratio of 28,30-bisnorhopane to norhopane; h35/h34, ratio of pentakishomohopane to tetrakishomohopane; cpi, carbon preference index. takes place in deposits containing both type ii and type iii kerogen, apparently with no significant differences in enrichment factors (disnar & sureau 1990). cuttings samples from this interval indicate the persistence of black laminated mudstones, suggesting that there were no significant changes in redox conditions in the sedimentary environment. the third possible factor, sedimentation rate, cannot be evaluated meaningfully on the basis of the available data. in general terms, the geochemical characteristics of the bo member conform to those of the remainder of the farsund formation as outlined above. however, a number of specific characteristics serve to geochemically differentiate the bo member from the remainder of the farsund formation (figs 14–16). firstly, n-alkane distributions may display a slightly increased abundance of nc15 and/or nc17, and in the c20–28 range, a slight predominance of even carbon numbered components is sometimes noted, for example in the edna-1 well. in the m/z 191 ion fragmentogram, the proportion of tricyclic triterpanes relative to pentacyclic triterpanes, shown by the ratio of c23 tricyclic triterpane (t23) to c30 hopane (h30), may be increased (fig. 14). the abundance of 28,30-bisnorhopane (h28) is often very high (fig. 16), and in samples of low to moderate thermal maturity, this compound may even dominate the m/z 191 fragmentogram. extended hopanes may be abundant. the c35 and, in some cases, the c33 homologues are slightly enriched, leading to c35/c34 homologue ratios close to unity or above. geochemical interpretation of the bo member a predominance of nc15 and nc17 is generally assumed to indicate algal organic matter (gelpi et al. 1970; tissot & welte 1984). an abundance of tricyclic triterpanes has been linked to the occurrence of tasmanites-type alginite (azevedo et al. 1992; revill et al. 1994), which is present in large proportions in the bo member (bojesenkoefoed 1988). the presence of 28,30-bisnorhopane is 428 400 425 450 475 500 525 550 0 200 800 400 1000 600 tmax (°c) h yd ro ge n in de x ( h i) total organic carbon (%) poor .2 .3 .7 .8.9.4 .6.5 0.1 1 2 3 7 8 94 65 20 30 40 50 10 2 0.1 4 6 8 2 2 4 6 8 4 6 8 2 1 10 100 excellent ex ce lle nt good g oo d fa ir po ors2 ( kg h yd ro ca rb on s/ to n r oc k) type i type ii type iii a b fig. 15. rock-eval/toc data for samples of the bo member in various sub-basins of the danish central graben. a: tmax vs. hydrogen index (hi); dashed lines indicate maturity evolution paths for kerogen types i, ii and iii. note the elongate distribution of data points, mainly reflecting variation in thermal maturity of the bo member in the different sub-basins. b: total organic carbon (toc) vs. s2 pyrolysis yield; note the overall organic richness and high pyrolysis yields. generally assumed to indicate highly anoxic environments, and its occurrence has also been linked to bacterial activity (katz & elrod 1983; williams 1984; moldowan et al. 1985; peters & moldowan 1993). a predominance of even-numbered n-alkanes in the c20–28 range is often observed in mildly hypersaline depositional environments (welte & waples 1973; nishimura & baker 1986; grimalt & albaiges 1987; bojesen-koefoed et al. 1997) but has also been recorded from a marine setting with little terrigenous input (kennicutt & brooks 1990). a high proportion of homohopanes is generally favoured by strongly reducing environments (peters & moldowan 1993), and enrichment in c33 and c35 homologues is often observed in carbonate and hypersaline environments (mello et al. 1988). hence, the biomarker distribution is indicative of anoxia, a minimal input of organic matter derived from higher land plants and perhaps of conditions of mild hypersalinity. as noted earlier, the vast majority of commercial as well as non-commercial petroleum accumulations in the danish north sea can, with a high degree of certainty, be genetically related to the farsund formation (and its stratigraphic equivalents). a common feature of almost all crude oils from the danish north sea is 429 10 20 30 40 50 60 7010 20 30 40 50 60 70 36 38 40 42 44 46 48 50 30 40 50 60 abundance abundance abundance abundance m/z 191 m/z 217 36 38 40 42 44 46 48 50 60504030 abundance abundance m/z 191 m/z 217 time (min.) time (min.)time (min.) time (min.) (a) (b) bo member farsund formation 17 ph pr 17 { { { { 14 15 {16 17 18 19 20 21 22 23 24 14 15 {16 17 18 19 20 21 22 23 24 4 1 2 3 5 6 7 8 9 10 { 11 12 13 1 2 3 5 6 7 8 9 10 { 12 13 { { 11{ { pr ph time (min.)time (min.) fig. 16. comparison of the geochemical characteristics of the bo member (a) and ‘background’ farsund formation mudstones (b). the uppermost pair of traces are gas chromatograms; the solid line connects n-alkanes with even numbers of carbon atoms, the dashed line connects n-alkanes with odd numbers of carbon atoms. note the dominance of even-numbered n-alkanes in the bo member. the two pairs of traces beneath illustrate biomarker data, ion fragmentograms m/z 191 and m/z 217. note the relative abundance of c23 tricyclic triterpanes (peak number 1) and 28,30-bisnorhopane (4) in the bo member mudstones. additional peaks: 2, ts; 3, tm; 5, norhopane; 6, c29-moretane; 7, hopane; 8, c30-moretane; 9–13, homohopanes; 14–17, c27-diasteranes; 18–21, c27 regular steranes; 22–24, c29 regular steranes. the presence of varying proportions of 28,30-bisnorhopane. this compound is thermally labile and its presence will to some extent be governed by maturity; indeed, the variation observed roughly parallels the maturity of the oils (unpublished geus data). since 28,30-bisnorhopane is only found in appreciable proportions in the sediments of the bo member, it may be assumed that this member has contributed to most of the hydrocarbon occurrences in the danish north sea. this further supports the idea that the bo member is a persistent feature in the danish central graben, and probably retains its identity in areas outside present well control, such as in the central portions of the feda and gertrud grabens and along the western flank of the central graben. furthermore, the occurrence of a large number of ‘immature’ oils, particularly in the southern part of the danish central graben, all carrying notable proportions of 28,30-bisnorhopane, suggests that the kerogen of the bo member and the upper farsund formation in general is able to generate and expel petroleum at low levels of thermal maturity. discussion upper jurassic organic-rich shales: current models the importance of upper jurassic source rocks in northwest europe has resulted in numerous studies focussing on the mode of accumulation and preservation of organic matter in this intra-cratonic setting (e.g. tyson et al. 1979; oschmann 1988; miller 1990; wignall 1991a). there is general agreement that a stratified water column with anoxic or dysoxic bottom waters, at least periodically, is indicated by the lithofacies, biofacies, palynofacies and geochemistry. rather less agreement has been reached, however, concerning the dominant mechanism(s) controlling such stratification and anoxia. tyson et al. (1979) suggested that stagnation and stratification in the kimmeridgian sea was influenced both by the regional palaeogeography and the overall high sea level; the complexity of the former inhibited circulation and open ocean transfer while the increased water depths over the shelf favoured stratification of bottom waters beneath wave-base. in some respects, this resembles the ‘silled basin model’ of demaison & moore (1980), using the quaternary of the black sea as a broad analogy; the onset of stratification on the late jurassic shelf was envisaged as a response to thermocline development, however, rather than to salinity stratification. the role of sea-level variation in controlling the regional extent of bottom-water anoxia has been emphasised (wignall 1991a; wignall & hallam 1991) while the link between cyclical variations in source rock development and climatic fluctuations has been discussed by a number of workers (e.g. oschmann 1988). alternative models have been presented to explain the regional development of a stratified water mass in northwest europe during the late jurassic – earliest cretaceous. oschmann (1988) suggested that the ‘silled basin’ model was inappropriate, given the regional extent of the organic-rich facies and the well-established connections to the tethyan and boreal seas. he argued for the seasonal development of extensive anoxia due to the southward migration of cold, oxygen-poor boreal waters during the summer in response to a northerly, wind-driven surface current. turnover of this stratified system probably occurred during the winter in all but the deepest submarine grabens. miller (1990) proposed a simple two-layer oceanographic model that is grossly the reverse of oschmann (1988): cold oxygenated boreal waters formed the surface layer flowing southward above warm saline bottom waters that originated in hypothetical shallow evaporative bays and flowed sluggishly northwards collecting in local depocentres and major rift axes. this model requires a sensitive balance between temperature and salinity of the two water masses since slight cooling of the boreal waters or decrease in the elevated salinities of the warmer southern waters would result in complete overturn. such a scenario, initiated by climatic or oceanographic shifts, was envisaged by miller (1990) for the regional destabilisation of the stratified system in the latest ryazanian (the basin ‘flushing’ of rawson & riley 1982). volgian–ryazanian ‘hot shales’ in discussion of the models proposed for the development of the ‘kimmeridge clay formation’ sea, it should be noted that such models are based largely on onshore data, particularly from the type area (see discussion by miller 1990). the facies are thus not directly comparable to those of the central graben, and indeed the depositional models are only strictly applicable to the period represented by the onshore section (i.e. kimmeridgian – middle volgian). furthermore, such models are designed to explain the long-term controls on the deposition and preservation of organic-rich shales that characterise the ‘kimmeridge clay formation’ in all 430 its stratigraphic guises throughout northwest europe. the ‘hot shales’ of the uppermost jurassic – lowermost cretaceous in the central graben of the north sea record a distinctive event within the background of organic-rich shale sedimentation, perhaps precipitated by the enhancement of one controlling factor or a coincidental combination of factors. the data presented here from the danish central graben are assessed below in the light of the current models summarised above. stratigraphic continuity in attempting to understand the origin of these organicrich shales, it is clearly important to establish the regional extent and degree of synchroneity of these deposits. black, organic-rich shales are a characteristic feature of the kimmeridgian–ryazanian of northwest europe and indeed farther afield (ager 1975; doré et al. 1985; klemme 1994). clearly, this represents a time during which burial and preservation of organic carbon was favoured on a global scale. it is also clear, however, that local factors such as structural configuration and sediment influx in addition to short-term global variables (e.g. sea level, climate) controlled the degree of development in any one location. in the north sea basin, the regional development of organic-rich shales in the kimmeridge clay formation (and stratigraphic equivalents) shows marked diachronism, although peak developments may be more widespread and biostratigraphically correlatable. doré et al. (1985), discussed at length the temporal and geographic distribution of ‘hot shales’ in the north sea region. these workers proposed that deposition of organic-rich ‘hot shales’ was most extensive, temporally, in the viking graben of the northern north sea where much of the kimmeridgian and volgian stages are represented by anaerobic, organicrich claystones in distal basin-centre locations. according to doré et al. (1985), such facies are not well-developed in the ryazanian succession of the viking graben, however, in contrast to the central graben farther south. as noted earlier, ‘hot shales’ occur at a number of levels in the upper jurassic of the north sea basin, the middle volgian – ryazanian succession under focus here forming the uppermost and best-developed example (price et al. 1993). in the norwegian sector of the danish basin, hot shales of kimmeridgian – early volgian age form a mappable unit defined as the tau formation (hamar et al. 1983; doré et al. 1985). as observed by rawson & riley (1982), the onset of anoxia in the kimmeridgian marked by the appearance of the tau formation organic-rich shales corresponds to the base of the oil shale facies in the eudoxus chronozone in the type kimmeridge clay formation of the wessex basin (tyson 1996). this event is marked by thin ‘hot’ log spikes within the farsund formation of the southern norwegian (doré et al. 1985) and danish sectors of the central graben (johannessen et al. 1996). according to wignall (1991b), this is one of the most important flooding events recorded in the onshore kimmeridge clay formation. from the outer moray firth through the central graben to the netherlands sector (fig. 1), the volgian– ryazanian succession is characterised, to a greater or lesser extent, by ‘hot’ organic-rich shales. on the basis of available published data, it appears that the onset of the development of this organic-rich facies may not have been wholly synchronous. in the uk sector of the central graben, donovan et al. (1993) ascribed an early volgian age to this event whereas a middle–late middle volgian age is likely in the norwegian and danish sectors (doré et al. 1985; dybkjær 1998). the same event has been referred to the late middle – late volgian in the northern part of the netherlands sector (g.f.w. herngreen, personal communication 1995). although these sources suggest a crude north-to-south younging of the initiation of ‘hot shale’ sedimentation, caution should be exercised. firstly, there is a scarcity of well-documented biostratigraphic information and secondly, watermass stratification may well have been regionally synchronous throughout the central graben, yet local factors (organic productivity, siliciclastic dilution, basin morphology etc) may have dictated when organic-rich mudstones began to be preserved at any one location. in any event, existing data suggest that there is considerable overlap in the age of the most organic-rich facies; true ‘hot shales’ accumulated from the latest volgian to the mid-ryazanian in most parts of the central graben and the moray firth. depositional and structural setting the data from the danish sector reinforce the regional interpretations of the upper jurassic mudstones in the central graben and onshore (tyson et al. 1979; miller 1990; tyson 1996). sedimentological core data, combined with palynological and geochemical data, attest to predominantly anoxic conditions on the sea floor; suboxic conditions may have prevailed periodically in certain locations. regional stratigraphic data from the danish central graben indicate that the mid-volgian – 431 ryazanian succession is sediment-starved relative to the kimmeridgian – mid-volgian (andsbjerg & dybkjær 2003, this volume). this period of reduced sediment supply to basinal settings has been correlated with an arid climatic episode between the early volgian and the late ryazanian (hallam 1985; wignall & ruffell 1990; hallam et al. 1991; ruffell & rawson 1994), an event which is also reflected in the upwards decrease in the terrigenous component of the kerogen of the farsund formation in the danish central graben (see above). with the exception of the easternmost portion of the danish central graben, where increased siliciclastic supply is thought to have suppressed the organic signal, the bo member is recognised throughout the danish central graben, even on late jurassic intrabasinal highs. position within the complex segmented central graben clearly influenced the degree of development of the ‘hot shales’ (i.e. thickness, organic richness, timing of initiation etc. – see earlier discussion and figs 4, 6). at its peak, however, watermass stratification and resultant bottom-water anoxia was sufficiently pervasive as to extend over most structural highs within the danish central graben. role of sea-level variation in assessing the possible contribution of sea-level variation to the development of a stratified watermass and bottom-water anoxia, it is important to recognise that the volgian–ryazanian succession records organic carbon burial and preservation at two different temporal scales. enhanced preservation of organic carbon began in the middle–late middle volgian in the danish sector and persisted until the early late ryazanian. the organicrich shales of the bo member, however, record a discrete short-term pulse, essentially restricted to the early ryazanian, that was superimposed on the long-term trend. referring, then, to published sea-level curves (hallam 1988; haq et al. 1988), it is apparent that the middle volgian – late ryazanian period was characterised, in broad terms, by a downward trend in sea level following the long-term sea-level rise that dominated most of the late jurassic and peaked in the early volgian. superimposed on this ‘second order’ fall in sea level, however, were a number of short-term sea-level events, as indicated by rawson & riley (1982, p. 2630) who referred to the middle – late volgian period as a “strongly regressive phase with occasional transgression”. one such minor transgressive event is recorded, for example, by the quasi-marine cinder bed which records a base kochi chronozone (early ryazanian) marine incursion into the largely non-marine wessex basin (rawson & riley 1982). an early ryazanian transgressive trend is also reflected in the regional onlap of ryazanian ‘hot shales’ onto the flanks of the central graben in the uk sector (gatliff et al. 1994), culminating in the deposition of upper ryazanian black laminated mudstones on kimmeridgian strata in eastern england (lott et al. 1986, 1989). it is apparent, therefore, that conditions favourable for the deposition and preservation of organic-rich sediments in the central graben from the middle–late middle volgian to the early late ryazanian coincided broadly with a long-term fall in sea level. in contrast, the pulse of highly efficient burial of organic carbon in the early ryazanian, represented by the bo member, may correlate with a short-term sea-level rise; this is compatible with the evidence presented here from the lower boundary of the bo member of a marked decrease in sediment supply to the deep basin at this time. conclusions it seems likely that the enhanced anoxia that developed in the central graben of the north sea in the latest jurassic and earliest cretaceous was the result of a coincidence of several factors. firstly, the late jurassic intracratonic seaway in the north sea region was prone to stratification and the development of oxygen deficiency in bottom waters, probably as a result of regional palaeogeographic and partially related oceanographic factors (tyson et al. 1979; doré 1991). a number of additional factors combined to create conditions particularly favourable for the accumulation of highly organic-rich sediments in the late–middle middle volgian – early late ryazanian. an arid climate resulted in reduced sediment influx and, hypothetically, the development of saline waters in fringing evaporative bays that may have sourced the central graben bottom waters, thus enhancing watermass stratification (miller 1990). a long-term ‘second order’ fall in sea level from the middle volgian to the late ryazanian, in combination with the complex tectonic topography of the north sea, may have contributed to the poor circulatory system. it appears likely, however, that the lower ryazanian ‘hot shales’, represented by the bo member in the danish sector, record a superimposed, short-term transgressive event (or events) that resulted in lateral expansion of the anoxic bottom waters onto the flanks of the central graben and over intra-basinal highs. 432 acknowledgements this study incorporates results from a series of energy research projects (efp 91, efp 92, efp 94), funded in part by the danish ministry of environment and energy. we thank jan andsbjerg, peter n. johannessen, jens jørgen møller and niels e. poulsen for advice and constructive criticism. the paper was greatly improved by the reviews of richard v. tyson and paul b. wignall. references ager, d.v. 1975: the jurassic world ocean (with special reference to the north atlantic). in: finstad, k.g. & selley, r.c. 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(eds): tectonic evolution of the north sea rifts. publication (international lithosphere program) 181, 1–36. new york: oxford university press. manuscript received 1 april 1996; revision accepted 1 november 1997. geological survey of denmark and greenland bulletin 38, 2017, 45-48 45 the onshore cretaceous–paleocene nuussuaq basin in west greenland (fig. 1) has long served as an analogue for offshore petroleum exploration. with the discovery of oil seeps on disko, nuussuaq, ubekendt ejland and svartenhuk halvø in the early 1990s, onshore exploration was also carried out. this eventually resulted in the gro#3 wildcat exploration well on western nuussuaq in 1996, which showed several intervals with hydrocarbons (christiansen et al. 1997). recent photogrammetric mapping of conspicuous marker horizons within the volcanic sequences of the basin shows that significant compressional structures may have developed in the latest paleocene on central nuus suuaq and northern disko that could be promising potential exploration targets. regional geological setting of the nuussuaq basin the nuussuaq basin, central west greenland, is a rift basin that developed during the cretaceous–paleocene in response to regional extension between greenland and canada. it is situated at the north-eastern edge of a complex system of rift basins and transfer systems that linked extension and sea-floor spreading in the labrador sea to the baffin bay (fig. 1). the basin was formed by two major phases of extension in the early cretaceous and late cretaceous, with an intervening quiescent period of thermal subsidence, when thick successions of source-prone mudstone were deposited regionally (see dam et al. 2009 for a detailed summary of the lithostratigraphy of the basin). significant volcanism beginning in the paleocene resulted in the deposition of a thick volcanic succession (fig. 2, electronic supplementary (es) figure: fig. es1; larsen et al. 2016). the supplementary material includes a summary of the complete volcanic and sedimentary stratigraphy. oakey & chalmers (2012) document significant changes in the kinematic evolution of the baffin bay and labrador sea during the latest paleocene–eocene (magnetic chrons c25n–c24n) that are related to the opening of the north atlantic ocean. based on seismic-stratigraphic interpretation constrained by wells, this time also marks the apparent onset of inversion in the offshore basins (gregersen & bidstrup 2008). inversion structures as potential petroleum exploration targets on nuussuaq and northern disko, onshore west greenland erik v. sørensen, john r. hopper, gunver k. pedersen, henrik nøhr-hansen, pierpaolo guarnieri, asger k. pedersen and flemming getreuer christiansen baffin island greenland eocene oc paleocene oc offshore sb onshore sb transitional crust palaeogene basalts buried volcanic rocks extensional fault thrust/reverse fault transform fault extinct spreading axis nuussuaq basin davis strait high figure 1 fig. 1. regional setting of the nuussuaq basin, simplified from oakey & chalmers (2012). oc: oceanic crust. sb: sedimentary basin. © 2017 geus. geological survey of denmark and greenland bulletin 38, 45–48. open access: www.geus.dk/publications/bull 4646 photogrammetric mapping of inversion structures during the danian and earliest selandian, large volumes of picritic lava were erupted in the southern part of the nuussuaq basin, forming the vaigat formation (e.g. larsen & pedersen 2009). the formation is divided into three main members (fig. 2) that primarily consist of greyish weathering, mg-rich, picritic rocks. however, intervals of brown to light-coloured, crustally contaminated siliceous basalts to magnesian andesites that make good marker horizons also occur throughout the succession. two marker horizons in the uppermost nujanáguit member (fig. 2) are regional in extent, easily mappable, and originally formed a sub-horizontal surface, referred to as the tunoqqu surface. photogrammetric mapping shows that the tunoqqu surface is now segmented into areas of different elevation and structural trends as a result of later tectonic deformation (sørensen 2011). this is most notable on nuussuaq where the western part is elevated and in part highly faulted. around the qunnilik valley, the surface has been uplifted and faulted into many small blocks by numerous faults, so that it now forms an asymmetric anticline with a steeper dipping western limb and a gently dipping eastern limb (fig. 3). measured vertical displacement on faults varies from a few metres to around 100 m, whereas the amplitude of the folding, measured as the elevation difference between the axial parts of the syncline and anticline amounts to around 900 m. the limbs of the anticline are coincident with two extensional faults that pre-date the tunoqqu surface, the kuugannguaq–qunnilik (k–q) and p faults of chalmers et al. (1999). the main fold axis appears to have an overall n–s trend (fig. 4), although in detail there may be local variations. the details, however, are difficult to resolve with the mapping technique, so some caution should be used when interpreting details. seismic data from the vaigat show evidence that underlying strata are also folded (marcussen et al. 2002). these too indicate a n–s axial trend in the folds. the exact timing of the inversion is difficult to resolve, but must post-date the deposition of the naujánguit mb. it is most likely a very late paleocene structure and thus formed at the same time as the onset of offshore inverdanian selandian thanetian c27 c26r c25r 60.2 ± 0.5 niaqussat mb nordfjord mb rinks dal mb 61.2 ± 0.4 ordlingassoq mb 61.3 ± 0.5 naujánguit mb 61.2 ± 0.5 sv ar te nh uk f m m ali gâ t f m va iga t f m anaanaa mb ma dykes 57.49 ± 1.40 ma 58.34 ± 0.40 gabbro sill 58.96 ± 0.51 figure 2 60 62 61 59 fig. 2. summary of the paleocene volcanic stratigraphy of the nuussuaq basin from larsen et al. (2016). the red line marks the stratigraphic position of the tunoqqu surface. more complete sedimentary and volcanic stratigraphy from published material is available as an electronic supplement (fig. es1). 54°w 55°w 53°w 70°n 70°30’n it p ik m k it k-q p as ma 20 km 1600 m 700 m 1200 m elevation of tunoqqu surface syncline anticline gro#3 q figure 3 disko hareøen nuussuaq uummannaq fjord vaigat major faults other faults inferred fault fig. 3. tunoqqu surface mapped by photogrammetry. as: asuk locality. ma: marraat locality. ik: ikorfat fault. it: itilli fault. k–q: kuugannguaq–qunnilik fault. faults ‘p’ and ‘m’ follow the nomenclature of chalmers et al. (1999). note that fault ‘m’ is inferred from gravity modelling, not surface geology. location of the gro#3 well is also shown. note that in the vaigat, folding with an overall n–s trend is indicated on seismic reflection data. 47 sion (gregersen & bidstrup 2008). guarnieri (2015) suggests an e–w-directed, compressional, palaeostress regime along west greenland during the latest paleocene that is consistent with the orientation of the structure. whether the inversion was a short-lived event or took place during a longer period of time is less clear from the present data. in any case the ne–sw-trending itilli fault, an important strike-slip fault active during the eocene, shows a leftlateral movement that seems to be incompatible with n– s-trending compressional folds on central nuussuaq and northern disko. for this reason the activity of the itilli fault likely post-dates the tectonic inversion, suggesting a short-lived period for the compressive event. distribution of potential source rocks hydrocarbon seeps have been mapped in the region and five distinct oil types have been identified (fig. 4; bojesenkoefoed et al. 1999). two oil types are particularly important for exploration: the marraat oil and the itilli oil. the source rock for the marraat oil was sampled in the gro#3 well within the marine, syn-volcanic eqalulik formation. the source rock for the itilli oil has not been sampled, but is interpreted to be of cenomanian–turonian age or older and have a wide distribution (bojesen-koefoed et al. 1999). although currently unproven, this interval is expected to be present in the lower itilli formation in the region (bojesen-koefoed et al. 1999). figure 4 shows inferred distribution of the two most important source-prone formations. the distribution of the itilli formation is regarded to be of regional extent, extending west and north-west into the davis strait and baffin bay. based on sediment thicknesses modelled by chalmers et al. (1999), it is suggested here that the lower itilli formation was probably sufficiently buried to have generated oil in large areas west and north-west of the ikorfat fault, although the timing of hydrocarbon generation is highly uncertain. the map is thus consistent with the broad distribution of the itilli oil type observed throughout the region. in contrast, the region where the eqalulik formation may have been sufficiently buried to generate oil is likely more limited. here it is suggested that the oil potential of the formation is restricted to the west of the kuugannguaq–qunnilik fault and thus is not likely to migrate to areas east of the fault, consistent with the lack of marraat oil observed in areas other than south-west nuussuaq. ? ? ? ? ? inferred outcrop itilli fm not preserved inferred outcrop poorly preserved eqalulik fm 54°w 70°n 71°n 71°n 54°w 70°n areas that may be source prone areas that may be source prone itilli oil seep marraat oil seep faultsfaults disko nuussuaqnuussuaq uummannaq fjord uummannaq fjord vaigat svartenhuk halvø ubekendt ejland ubekendt ejland upernivik ø hareøen disko nuussuaqnuussuaq vaigat svartenhuk halvø upernivik ø hareøen figure 4 a nt ic lin e ax is a nt ic lin e ax is fig. 4. inferred distribution maps of the itilli and eqalulik formations. black lines are faults (see fig. 3). the overall trend of fold axis of the anticline mapped is shown by the red line. the maps are based on the known distributions from onshore outcrops and from offshore seismic data that indicate the presence of significant cretaceous–paleocene sedimentary strata (marcussen et al. 2002). also shown are the locations of marraat and itilli oil seeps and stains. it is notable that the marraat oil is concentrated only on nuussuaq between the qunnilik and itilli faults whereas the itilli oil is known regionally. the distribution of areas where there may be oilprone intervals is shown by the grey shading. this is highly speculative and based on the locations of oil seeps and outcrops. along southern nuussuaq and northern disko, the lower itilli formation is absent east of the k–q fault, but recent samples collected suggest it is present near the ikorfat fault along northern nuussuaq. 4848 conclusions oil and gas shows in cores, along with numerous oil seeps, attest to the fact there is a working petroleum system in the region. previous exploration on western nuussuaq where seeps are most abundant failed to identify viable traps and early exploration was therefore abandoned after drilling of the gro#3 well. however, the structural anticline defined by the tunoqqu surface covering an area of ~250 km2 on central-west nuussuaq suggests that large structures could well be present in the region and in the underlying sedimentary rocks. previous play concepts in the region generally assume that the main source rocks are to the west or south-west of the main oil-seep areas, i.e., in the main deep marine depocentres. the areas farther east have been considered to be less prospective, since any oil would have to migrate longer distances and bypass the faulted area around the k–q fault. the presence of oil seeps at asuk far to the east of this fault is thus enigmatic, raising questions about the source rock. here, we suggest a new lead concept and propose that the central-west nuussuaq and uummannaq fjord areas also hold potentially mature source rocks. this would imply that the region west of the ikorfat fault is prospective on nuussuaq and possibly also in vaigat and on northern disko. the migration path of the oils found at asuk could very well be from the north, rather than the west. acknowledgements this contribution is partly the result of a project funded by the ministry of mineral resources, greenland. discussions and work with niels h. schovsbo, thomas f. kokfelt, morten l. hjuler, christian knudsen, peter johannessen, jim a. chalmers, lotte m. larsen, jørgen bojesenkoefoed, ulrik gregersen, peter japsen, jens-jørgen møller, and nina skaarup were helpful during the course of the project. references bojesen-koefoed, j.a., christiansen, f.g., nytoft, h.p. & pedersen, a.k. 1999: oil seepage onshore west greenland: evidence of multiple source rocks and oil mixing. in: fleet, a.j. & boldy, s.a.r. (eds): petroleum geology of northwest europe: proceedings of the 5th conference, 305–314. london: geological society. bojesen-koefoed, j.a., bidstrup, t., christiansen, f.g., dalhoff, f., nytoft, h.p., nøhr-hansen, h., pedersen, a.k. & sønderholm, m. 2007: petroleum seepages at asuk, disko, west greenland: implications for regional petroleum exploration. journal of petroleum geology 30, 219–236. chalmers, j.a., marcussen, c. & pedersen, a.k. 1999: new insight into the structure of the nuussuaq basin, central west greenland. marine and petroleum geology 16, 197–224. christiansen, f.g., boesen, a., dalhoff, f., pedersen, a.k., pedersen, g.k., riisager, p., & zinck-jørgensen, k. 1997: petroleum geological activities onshore west greenland in 1996, and drilling of a deep exploration well. geology survey of greenland bulletin 176, 17–23. clarke, d.b. & pedersen, a.k. 1976: tertiary volcanic province of west greenland. in: escher, a. & watt, w.s. (eds): geology of greenland, 365–385. copenhagen: grønlands geologiske undersøgelse. dam, g., pedersen, g.k., sønderholm, m., midtgaard, h., larsen, l.m., nøhr-hansen, h. & pedersen, a.k. 2009: lithostratigraphy of the cretaceous–paleocene nuussuaq group, nuussuaq basin, west greenland. geological survey of denmark and greenland bulletin 19, 171 pp. gregersen, u. & bidstrup, t. 2008: structures and hydrocarbon prospectivity in the northern davis strait area, offshore west greenland. petroleum geoscience 14, 151–166, http://dx.doi.org/10.1144/1354079308-752 guarnieri, p. 2015: pre-break-up palaeostress state along the east greenland margin. journal of the geological society 172, 727–739, http:// dx.doi.org/10.1144/jgs2015-053 larsen, l.m. & pedersen, a.k. 2009: petrology of the paleocene picrites and flood basalts on disko and nuussuaq, west greenland. journal of petrology 50, 1667–1711. larsen, l.m., pedersen, a.k, tegner, c., duncan, r.a., hald, n. & larsen, j.g. 2016: age of tertiary volcanic rocks on the west greenland continental margin: volcanic evolution and event correlation to other parts of the north atlantic igneous province. geological magazine 153, 487–511, http://dx.doi.org/10.1017/s0016756815000515 marcussen, c., skaarup, n. & chalmers, j.a. 2002: efp project nuussuaqseis 2000: structure and hydrocarbon potential of the nuussuaq basin: acquisition and interpretation of high resolution multichannel seismic data. danmarks og grønlands geologiske undersøgelse rapport 2002/33, 63 pp. oakey, g.n. & chalmers, j.a. 2012: a new model for the paleogene motion of greenland relative to north america: plate reconstructions of the davis strait and nares strait regions between canada and greenland. journal of geophysical research 117, http://dx.doi. org/10.1029/2011jb008942 sørensen, e.v. 2011: implementation of digital multi-model photogrammetry for building of 3d-models and interpretation of the geological and tectonic evolution of the nuussuaq basin, 204 pp. copenhagen: unpublished phd thesis. authors’ addresses e.v.s, j.r.h., g.k.p & p.g., geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: evs@geus.dk. a.k.p., natural history museum of denmark, øster voldgade 5-7, dk-1350 copenhagen k, denmark. geological survey of denmark and greenland bulletin 33, 2015, 37-40 37© 2015 geus. geological survey of denmark and greenland bulletin 33, 37–40. open access: www.geus.dk/publications/bull assessment of the mineral raw material potential in denmark – methods and major findings jakob kløve keiding, per kalvig, claus ditlefsen, steen lomholt and peter roll jakobsen aggregates and other mineral raw materials are important prerequisites for the continual development of the infrastructure and economic growth of a country. th e production of these raw materials in denmark amounted to c. 4.5 m3 per capita in 2012, which was 57% higher than the average in eu and efta countries (uepg 2014). in this perspective, it is essential to locate and assess the danish mineral resources in order to plan future exploitation, especially in densely populated regions where both spatial competition for landuse and demands for raw materials are high. here we present the methods used in a recent resource evaluation that for the fi rst time includes danish resources both on land and at sea and summarises some of the main fi ndings of this analysis. th e importance of availability of mineral raw materials such as sand, gravel, clay and limestone for a society cannot be overestimated: they are prerequisites for the development of the infrastructure and a wide range of products for the industrial sector. th e supply of these raw materials is the result of complex value chains, which in turn are based on indicated (see below) and exploited mineral resources from which the raw materials can be processed and turned into commercial materials and products. globally, sand and gravel account for the largest volume of solid material extracted, with c. 15–25 billion m3 excavated annually, and demand increasing rapidly (geas 2014). th e danish mineral industry includes exploitation of loose aggregate, limestone and chalk, diff erent types of clay, salt and granite, and amounted to c. 35 million m3 in 2012 of which aggregate alone accounted for c. 90% (statistics denmark 2012a, b). according to a recent forecast, the demand for sand and gravel in 2036 will be c. 40 million m3 (regionernes videncenter for miljø og ressourcer 2014). th e greater part of the sand and gravel excavated in denmark is used as aggregate and fi ller for concrete and other construction purposes. limestone is mainly used for cement production as industrial fi ller and as soil conditioner. th e clay industry is fairly diversifi ed and uses clay for bricks, insulation stones, membranes and special products such as absorbents and ion-exchange media. salt is mainly used for de-icing of roads and in the food and chemical industries. about 35 000 persons were employed in the primary and secondary sectors of this part of the danish industry in 2013 with a turnover of about 3500 million dkk (statistics denmark 2012a, b). exploitation of mineral resources in denmark takes place at designated sites and the administration is controlled by the danish nature agency for the marine resources and the danish regions for the resources on land. th ere are about 630 onshore sites for excavation of sand, gravel, clay, limestone and granite, and about 100 off shore dredging sites for sand and gravel. are there suffi cient mineral resources in denmark to cover the future demand? sometimes it is postulated that denmark is endowed with abundant sand and gravel deposits, and aggregates are thus oft en considered an essentially unlimited resource that will be available for exploitation in the foreseeable future. th is is an oversimplifi ed assumption because (1) aggregates cover a diverse group of commodities and not all qualities and types are equally abundant, (2) the distribution of mineral raw materials refl ects the local geology and aggregates are geographically unevenly distributed and oft en have to be transported over considerable distances from the excavation site to the end-use location, and (3) there are many diff erent and oft en competing landuse interests particularly in a densely populated country like denmark; exploitation of raw materials does not necessarily have the highest priority. an assessment of the aggregate resources in terms of their size, type and location – in conjunction with an evaluation of possible areas of confl ict with other interests – is crucial for both future resource and landuse planning. on this background the center for minerals and materials (mima) conducted the fi rst comprehensive danish raw material resource assessment comprising both onshore and off shore areas. additional work on mapping the value chains related to the danish raw material sector is in progress. methods th is assessment covers the following resources: aggregates (sand, gravel and pebble); limestone and chalk; clay (clay for bricks and tiles; bentonitic clay; diatomite); salt and granite. aggregates cover a large and diverse group of materials that can be divided into diff erent quality classes according 3838 to their composition. traditionally, quality classes applied to marine and terrestrial resources are diff erent. in order to compare the marine and terrestrial inventories this assessment uses the classifi cation used for exploitation of marine deposits. it discriminates between six commercial classes (fig. 1): sand 0, sand 1, gravel 2, pebble 3, filler sand 4 and sand x (unspecifi ed quality). it is beyond the scope of this short paper to describe the defi nition of the various classes and qualities in detail but they refl ect diff erence in grain size, grain size distribution, as well as an assessment of the petrographic composition (i.e. the content of reactive, unstable and porous mineral grains). for further details see ditlefsen et al. (2015). th is survey is based on data available to the geological survey of denmark and greenland (geus) as per october 2014. th e evaluation of land-based resources is based on previous investigations, including mapping of raw materials reported to or conducted by geus, as well as new resource assessments conducted by regional authorities within the present extraction areas. borehole data and geological maps were also used. for the fi rst time we also incorporated data from geophysical surveys conducted for hydrogeological mapping. th is allowed us to screen large parts of the terrestrial area for potential undiscovered raw material resources. th e resources addressed in this paper are classifi ed by reliability of the resource data and are divided into three classes, namely measured, indicated and inferred. th e criteria for assigning the resources to the diff erent categories were described by larsen (1994) and ditlefsen et al. (2015). in essence, the measured resources encompass raw materials that are mapped in detail and have well-known quality parameters. th eir characterisation thus has a high degree of certainty (± 20%). indicated resources are oft en based on shallow geophysical surveys and a small number of boreholes and have a lower reliability. inferred resources are based on geophysical data only, or on relatively sparse geological data, regional-scale mapping, etc. to compile the nation-wide assessment, data from the communes (98 in total) and from 41 marine study areas were amalgamated. fig. 1. map of denmark showing the location and areal extent of the different quality classes of aggregate mineral resources used in this assessment. the map illustrates the combined resources of measured and indicated resources. 50 km 10°e 55°n 57°n 55°n 57°n sweden germany sand 0 sand 1 gravel 2 pebble 3 f i l ler sand 4 sand x, unspeci f ied qua l i ty german waters swedish waters norwegian waters bornholm jylland copenhagen 15°e 55°n based on an assessment of the profitable excavation depth, the compilation assesses the volume of raw materials that is available within the top 25 m below surface, except for salt, which is exploited from deeper-seated salt diapirs and therefore considered to a depth of 500 m below terrain surface. based on the same reasoning, we excluded areas with more than 5 m thick layers of soil or other material (overburden) covering the soughtaft er resource. when evaluating areas with limestone and chalk, only deposits with less than 10  m of over39 burden were considered. data for aggregates were compiled for both marine and terrestrial areas, whereas the remaining raw material groups were only considered on land since at sea they occur under a thick sedimentary cover and are therefore of no interest for exploitation. all areas with resources were screened with respect to commonly occurring confl icting interests (here referred to as ‘no-go’ fi lters). th is assessment thus specifi es both gross volumes, comprising the total volume of the resources and net volumes, taking into account the eff ect of ‘no-go’ fi lters. th e latter are calculated by subtracting volumes underlying areas that are claimed for other landuse from the total mineral resource fi gures. th e ‘no-go’ fi lter areas include nature protection areas (natura 2000 areas, historical monuments, state-protected areas, etc.), major built-up and paved areas, and marine waters less than 6 m deep where exploitation is prohibited according to present-day jurisdiction, but areas less than 0.01 km2 in size are excluded. since this is not a full list of areal restrictions, the available area for mineral extraction as well as the extractable net volumes will be smaller due to other local pre-existing priorities. results and conclusions th e mineral assessment presented here covers about 40% of the danish marine and terrestrial areas. possible resource occurrences have been indicated in about 10% of the surveyed areas. th e calculated gross and net mineral resources are summarised in table 1, and fig. 1 shows the geographical distribution of the various resource types and qualities of aggregates, the largest commodity group apart from salt. salt is the dominant resource in terms of volume for all three classes and comprises 89% of the measured net resource volumes (table 1). salt, however, is a special raw material that table 1. summary (in million cubic metres) of calculated gross and net mineral resources in denmark aggregates granite clay chalk and limestone salt total gross resources measured 4603 38 46 293 34 669 39 649 indicated 10 747 0 68 89 809 230 820 134 inferred 80 064 3175 1190 9192 489 671 583 292 net resources measured 4073 38 46 293 34 669 39 119 indicated 9621 0 66 89 809 230 819 006 inferred 68 888 2020 1029 7048 489 671 568 656 sum 95 414 3213 1304 9574 1 333 570 1 443 075 sum 82 582 2058 1141 7430 1 333 570 1 426 781 table 2. calculated net volumes (million cubic metres) of aggregate resources in denmark * unspecified quality. sand 0 sand 1 gravel 2 pebble 3 filler sand 4 sand x total marine resources measured 184 2273 207 77 111 2852 indicated 1214 6334 739 135 94 8516 inferred 497 1183 106 13 418 2217 terrestrial resources measured 191 656 18 191 165 1221 indicated 259 554 10 178 104 1105 inferred 43 96 0 13 66 519 66 671 * total assessment measured 184 2466 863 95 302 165 4073 indicated 1214 6593 1293 145 272 104 9621 inferred 497 1226 202 13 431 66 519 68 888 total sum 1895 10 283 2358 253 1005 66 788 82 582 sum (marine) 1895 9790 1052 225 623 13 585 sum (land) 493 1306 28 382 66 788 68 997 4040 diff ers from the other mineral resources in several ways: (1) a fundamentally diff erent extraction method is used, (2) it is only found at great depths and (3) and it is regulated by the danish subsoil act (undergrundsloven) in contrast to the other resources assessed here that are regulated by the raw materials act (råstofl oven). on a national scale, the measured resources of aggregates account for c. 10% of the registered mineral resources, followed by chalk and limestone (c. 1%; whereas the volumes of clay and granite account for less than 1%). for all the fi ve types of raw material considered, the volume of the inferred resource class is enormous (table 1). however, only minor fractions of these volumes will be available for exploitation due to confl icting landuse interests. aggregates mainly come from either glaciofl uvial deposits or from lateand postglacial marine deposits. marine deposits are more mature than glaciofl uvial deposits and usually contain smaller amounts of reactive grains that are harmful to constructions. in addition, older deltaic and fl uvial deposits of pure quartz sand are found in central jylland and on and around the island of bornholm. granites in denmark are confi ned to bornholm where they are part of the precambrian basement that is exposed or just covered by thin quaternary deposits on the northern two thirds of the island. clay is mainly of glaciolacustrine or marine origin, while limestone and chalk are marine formations of cretaceous and early cenozoic age. th e salt resources in denmark are permian in age and related to evaporates from the zechstein basin. most salt deposits are found at more than 3 km depth, and are of no economic interest. locally, however, the salt has been mobilised and is found in shallower diapirs at depths of 200–300 m. sand x (unspecifi ed quality) is the largest class in the aggregate group and comprises c. 81% of the total volume of all aggregates (table 2). th e resources in this class are currently based only on geophysical screening and require a more thorough characterisation with respect to grain size, grain-size distribution, petrographic composition as well as better constraints on the spatial distribution. however, this class off ers unique possibilities to focus future exploration for aggregates in relevant areas. sand 1, gravel 2, pebble 3 and filler sand 4 are the present commercial classes. th ese qualities are not equally distributed, hence about 95% of the sand 1 resources are located off shore, and about 88% of the pebble 3 resources are found onshore (measured resources, table 2). within the onshore areas, the sand and gravel resources are unevenly distributed, for example, very limited proven gravel resources are indicated close to the city of copenhagen. th ese observations are important for long-term planning of how to exploit the known resources. th ere is an up to 19% diff erence between distribution of gross and net deposits for the various raw material qualities when looking at measured and indicated resource classes. th e largest changes are seen for pebble 3 (19%), filler sand 4 (18%), sand 1 (10%) and sand 0 (7%). th e eff ect of the ‘no-go’ fi lters is generally rather limited for most of the commodities and raw material classes due to a few very large deposits. th is is in particular the case when including the inferred resources, which overshadows localscale restrictions, and thus in places the fi lters have a major impact on the resulting resource fi gures. also, a number of other constraints may be considered prior to designation of a resource resulting in substantial lower net resource fi gures. th is survey shows that denmark is endowed with enormous mineral resources of aggregates, clay and special clay, chalk and limestone, and salt and granite. however, to what extent these resources will be available as raw materials for the mineral industry is uncertain since it depends on local landuse and restrictions. finally, this assessment shows that both off and onshore aggregate deposits will have to be considered in order to secure the supply of all quality classes. references ditlefsen, c.b., lomholt, s., skar, s., jakobsen, p.r., kallesøe, a.j., keiding, j.k. & kalvig, p. 2015: danske mineralske råstofressourcer. kvantitativ analyse baseret på geologiske og geofysiske data. mima report 2015/1, 61 pp. copenhagen: geological survey of denmark and greenland. geas (global environmental alert service) 2014: sand, rarer than one thinks. www.unep.org/pdf/unep_geas_march_2014.pdf larsen, b. 1994: material sammensætningen i submarine råstoff orekomster – et metodestudium. dgu kunderapport 91, 51 pp. københavn: danmarks geologiske undersøgelse. regionernes videncenter for miljø og ressourcer 2014: fremskrivning af råstoff orbruget for 2013–2036. region syddanmark. råstoff er 4, 27 pp. statistics denmark 2012a: statistikbank – råstoff er indvundet fra havet. http://www.statistikbanken.dk/statbank5a/selectvarval/define. asp?maintable=rst3&planguage=0 statistics denmark 2012b: statistikbank – råstoff er indvundet på land. http://www.statistikbanken.dk/statbank5a/selectvarval/define. asp?maintable=rst01&planguage=0 uepg (union européenne des producteurs de granulats) 2014: european aggregates association. a sustainable industry for a sustainable europe. annual review 2013-2014. http://www.uepg.eu/uploads/ modules/publications/uepg-ar2013-2014_v28.pdf authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark; e-mail: jkk@geus.dk www.unep.org/pdf/unep_geas_march_2014.pdf http://www.statistikbanken.dk/statbank5a/selectvarval/define.asp?maintable=rst3&planguage=0 http://www.statistikbanken.dk/statbank5a/selectvarval/define.asp?maintable=rst01&planguage=0 http://www.uepg.eu/uploads/modules/publications/uepg-ar2013-2014_v28.pdf << /ascii85encodepages false /allowtransparency false /autopositionepsfiles true /autorotatepages /pagebypage /binding /left /calgrayprofile (dot gain 15%) /calrgbprofile (srgb iec61966-2.1) /calcmykprofile (coated fogra27 \050iso 12647-2:2004\051) /srgbprofile (srgb iec61966-2.1) /cannotembedfontpolicy /warning /compatibilitylevel 1.7 /compressobjects /off /compresspages true /convertimagestoindexed true /passthroughjpegimages false 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/romandefault /preserveediting true /untaggedcmykhandling /leaveuntagged /untaggedrgbhandling /leaveuntagged /usedocumentbleed false >> ] /syntheticboldness 1.000000 >> setdistillerparams << /hwresolution [2400 2400] /pagesize [595.276 841.890] >> setpagedevice geological survey of denmark and greenland bulletin 33, 2015, 57-60 57 greenland ice sheet melt area from modis (2000–2014) robert s. fausto, dirk van as, jens a. antoft, jason e. box, william colgan and the promice project team* th e greenland ice sheet is an excellent observatory for global climate change. meltwater from the 1.8 million km2 large ice sheet infl uences oceanic temperature and salinity, nutrient fl uxes and global sea level (ipcc 2013). surface refl ectivity is a key driver of surface melt rates (box et al. 2012). mapping of diff erent ice-sheet surface types provides a clear indicator of where changes in ice-sheet surface refl ectivity are most prominent. here, we present an updated version of a surface classifi cation algorithm that utilises nasa’s moderateresolution imaging spectroradiometer (modis) sensor on the terra satellite to systematically monitor ice-sheet surface melt (fausto et al. 2007). our aim is to determine the areal extent of three surface types over the 2000–2014 period: glacier ice, melting snow (including percolation areas) and dry snow (cuff ey & paterson 2010). monthly 1 km2 resolution surface-type grids can be downloaded via the cryoclim internet portal (www.cryoclim.net). in this report, we briefl y describe the updated classifi cation algorithm, validation of surface types and inter-annual variability in surface types. classification algorithm th e algorithm uses normalised thresholds (th ) from calibrated radiances (mod021km) between the near-infrared band 5 (1230–1250 nm) and the visible band 10 (483– 493  nm). th is updated classifi cation improves on fausto et al. (2007) by implementing new surface type thresholds: th dry snow≤0.86, 0.86<th melting snow<0.94 and th glacier ice≥0.94. data gaps can be caused by cloud cover, which is identifi ed from the mod35_l2. th e ratio between band 1 (620–670 nm, red visual) and band 7 (2105–2155 nm, shortwave infrared) as well as the magnitude of band 1 are used to identify non-glacierised terrain (fausto et al. 2007). daily classifi cation scenes are aggregated to yield a monthly greenland surface type (gst) product. for the gst product, pixels are classifi ed according to their most frequently occurring (cloud-free) colour during a month. th e maximum melt extent (gstmax) is defi ned by a given pixel classifi ed as melting or bare ice. th e pixel class must occur at least one day per month. th e gstmax product thus contains information about brief melt events. conversely, a minimum melt extent (gstmin) is calculated from pixels classifi ed as dry snow at least once per month. pixels from areas not cov* signe b. andersen, andreas p. ahlstrøm, morten l. andersen, michele citterio, charalampos charalampidis, karen edelvang, konstanze haubner, signe h. larsen, martin veicherts and anker weidick 22 50 250 km 2250 3000 2750 2500 1000 1500 1750 20 00 25 00 22 50 17 50 1500 1 2 5 0 2000 qas nuk kan thu upe kpc sco tas fig. 1. greenland map showing the locations of promice automatic weather stations. each dot represents two or three stations. elevation contours in metres. © 2015 geus. geological survey of denmark and greenland bulletin 33, 57–60. open access: www.geus.dk/publications/bull 5858 ered by ice are excluded from greenland melt area calculations employing the melting snow and glacier ice types in the gst, gstmax, and gstmin products (fausto et al. 2007). validation th e promice automatic weather station (aws) network currently consists of eight regions each with two or three stations at a variety of low elevations on the greenland ice sheet (fig. 1; ahlstrøm et al. 2008; citterio et al. 2015). most stations are located in the ablation area, and are thus transitioning from snow-covered to bare ice surfaces through the melt season. each station records a suite of meteorological and glaciological measurements, including ice temperature to 10 m depth and surface height changes due to accumulation or ablation (citterio et al. 2015). to validate the three surface types we classify, we use daily mean surface albedo (α), the ratio between incoming and outgoing solar radiation and surface temperature (t) data for the 2008–2014 period from the aws. based on α and t, we defi ne three classifi cations for aws sites: dry snow (α > 0.7 or t < –7°c), melting snow (0.7 > α > 0.55), and glacier ice (α < 0.55; cuff ey & paterson 2010). employing the t criterion acknowledges the infl uence of diurnal cycles at the beginning of the melt season. as a validation example, the in situ albedo and nearest-pixel classifi cation at kan_l in 2009 are presented in figs 2 and 3, including a visual comparison with the passive microwave melt area product (pmp) by mote (2007). th e kan_l station, located c. 10 km from the ice sheet margin at 680 m elevation, transitions through all three surface types during the melt season, from dry snow to melting snow to glacier ice. relative to the 2000–2014 period, the 2009 surface melt was normal in west greenland, with maximum melt areal extent in august. at kan_l, the surface melted from may to august, with a daily mean albedo generally between 0.5 and 0.6 (fig. 2). th e algorithm accuracy for the kan_l site may be assessed by an error matrix (table 1). th e diagonal represents successful classifi cations, the total number represents all classifi cations and the ratio between the sum of the diagonal and total is the accuracy. th e algorithm yields 79% successful classifi cations at kan_l, with an overall accuracy of 71%. th e classifi cation algorithm performs best in the south and worst in the north, with accuracies of 87% (nuk_l) and 61% (kpc_u), respectively. figure 2 illustrates changes in surface type during summer 2009, between 15 may and 14 september, according to the aws data; all but two classifi cations were successful. results and discussion th e melt area from this algorithm and the pmp of mote (2007), illustrated in fig. 3 for 12 july 2012, are consistent with the reported melt area by nghiem et al. (2012), who documented that 98.6% of the ice-sheet surface had melting. th e gst also demonstrates close visual correspondence with pmp for the 2000–2014 modis period (fig. 4). in fig. 4 we have plotted the yearly maximum values of the gst, gstmax and gstmin products, as well as the pmp maximum extent of greenland melt area. th e increasing trends of gstmax and gstmin indicate a rising frequency of melt events and increasing summer melt, which is corroborated by the pmp which is comparable with gstmax. th e trend for pmp between 1979 and 2000 and 2000 and 2012 are almost identical making the pmp and gstmax trends comparable. overall, an expansion of the melt area to higher elevations is apparent (fig. 4). fausto et al. (2007) suggested that a sub-monthly gst product is non-optimal, because missing data due to cloud cover is the primary problem in determining the melt area. when trying to characterise all of greenland, hall et al. (2012) also found clear-sky, day-count problems, and also suggested that a sub-monthly product would have signifi cant uncertainty. however, uncertainties associated with the difa lb ed o cl as si fic at io n 3 2 1 1 0.8 0.6 0.4 0.2 0 91 111 131 151 171 191 211 231 251 271 91 111 131 151 171 191 211 231 251 271 day of year day of year fig. 2. daily gst classification for 2009 of the kan_l pixel and albedo measured at the kan_l automatic weather station. 1: glacier ice. 2: melting snow. 3: dry snow. table 1. error matrix for the assessment of kan_l gst\aws* glacier ice melting snow dry snow total glacier ice 32 14 0 46 melting snow 3 33 1 37 dry snow 0 5 21 26 * gst: greenland surface type aws: automatic weather station 59 ferent surface types are assessed with the number of observations and standard deviation for each cloud-free pixel of the gst product (fausto et al. 2007). th e modis data have the advantages of high spatial resolution (1 km2), pan-ice sheet coverage and quasi-daily temporal coverage, while the footprints of the in situ measurements are small. th e aws surface type classifi cations are therefore not an ideal ground truth for the surface classifi cation. furthermore, whereas both gst and pmp melt area products can give daily results, the pmp surface microwave emittance originates not only from the surface but the top metre of the snow or fi rn, and is infl uenced by the water content in the snow during the previous days (mote 2007). modis classifi cation is sensitive to cloud cover, but the spatial resolution of pmp is 625 times coarser than gst. during the melt period, exposed glacier ice in the ablation zone can have sub-zero temperatures. such areas are included in the melting area, because the algorithm only makes use of the visual and near-infrared spectrum. hence the melt area that we map might be more representative of the cumulative melt area during the melt period. however, if exposed, glacier ice in the ablation zone is covered by snow it will be mapped as non-melting areas. an august anomaly in monthly gst is evident during the 2010–2014 period. all august images indicate a noisy melting snow classifi cation in the northern ice sheet (not shown), which is most likely due to false classifi cation. however, anomalous, high concentrations of dust or reddish material have been observed on the ice sheet during recent late summers (dumont et al. 2014). increasing dust concentrations are problematic for the fi xed threshold algorithm we employ, because of enhanced absorption in near infrared wavelengths. despite this possible biased source, an increasing trend in the melt area for the modis and pmp periods (fig. 4) is consistent with increasing greenland mass loss due to surface processes (tedesco et al. 2013). both independent, remotely sensed observations (hall et al. 2012) and in situ observations (mcgrath et al. 2013) show that the greenland melt area is expanding to higher elevations. fig. 4. yearly maximum melt area values and trends according to greenland surface type (gst), maximum melt extent (gstmax), minimum melt extent (gstmin) and passive microwave melt area product (pmp). 60°w 30°w 60°n 75°n a b no data/clouds melting snow dry snow glacier ice surface melting non-melting areas fig. 3. melt area on the greenland ice sheet for 12 july 2012 a: passive microwave melt area product (pmp). b: greenland surface type classification. m e lt a re a ( k m 2 ) 1800 1600 1400 1200 1000 800 600 400 200 0 1979 1986 1993 2000 2007 2014 time (year) gst gstmax gstmin pmp y = 9.0x − 17533 y = 35.3x − 69982 y = 3.3x − 6459 y = 22.1x − 43255 6060 conclusions th e modis data can yield daily, automated classifi cations of the greenland ice sheet into bare ice, melting and dry snow areas. validation indicates that the surface classes are useful as ice-sheet climate indicators. th e surface-type products are complementary to existing ice-surface temperature (hall et al. 2012) and melt-area (mote 2007) products. acknowledgements th e programme for monitoring of the greenland ice sheet (promice) is funded by the geological survey of denmark and greenland (geus) and the danish ministry of climate, energy and building under danish cooperation for environment in the arctic (dancea), and is conducted in collaboration with the national space institute (dtu space) and asiaq (greenland survey). th e nuk and kan stations were/are (co-)funded by the greenland climate research centre (gcrc) and the greenland analogue project (gap), respectively. th anks to t. mote for making the passive microwave product (pmp) available. th is study was funded by dk esa-prodex under the cryoclim project. references ahlstrøm, a.p. et al. 2008: a new programme for monitoring the mass loss of the greenland ice sheet. geological survey of denmark and greenland bulletin 15, 61–64. box, j.e., fettweis, x., stroeve, j.c., tedesco, m., hall, d.k., steff en, k. 2012: greenland ice sheet albedo feedback: thermodynamics and atmospheric drivers. th e cryosphere 6, 821–839. citterio, m., robert s. fausto, r.s., ahlstrøm, a.p., andersen, s.b., van as, d., charalampidis, c. & veicherts, m. 2015: automatic weather stations for basic and applied glaciology. geological survey of denmark and greenland bulletin 33, 69–72. cuff ey, k.m. & paterson, w.s.b. 2010: th e physics of glaciers, 693 pp. burlington: butterworth-heinemann/elsevier. dumont, m. et al. 2014: contribution of light-absorbing impurities in snow to greenland’s darkening since 2009. nature geoscience 7, 509– 512. fausto, r.s., mayer, c. & ahlstrøm, a.p. 2007: satellite-derived surface type and melt area of the greenland ice sheet using modis data from 2000 to 2005. annals of glaciology 46, 35–42. hall, d.k., comiso, j.c., digirolamo, n.e., shuman, c.a., key, j.r. & koenig, l.s. 2012: a satellite-derived climate-quality data record of the clear-sky surface temperature of the greenland ice sheet. journal of climate 25, 4785–4798. ipcc 2013: climate change 2013: th e physical science basis. working group i contribution to the fift h assessment report of the intergovernmental panel on climate change, 1535 pp. cambridge university press. mcgrath, d., colgan, w., bayou, n., muto, a. & steff en, k. 2013: recent warming at summit, greenland: global context and implications. geophysical research letters 40, 2091–2096. mote, t.l. 2007: greenland surface melt trends 1973–2007: evidence of a large increase in 2007. geophysical research letters 34, l22507. nghiem, s.v., hall, d.k., mote, t.l., tedesco, m., albert, m.r., keegan, k., shuman, c.a., digirolamo, n.e. & neumann, g. 2012: th e extreme melt across the greenland ice sheet in 2012. geophysical research letters 39, l20502. tedesco, m., fettweis, x., mote, t., wahr, j., alexander, p., box, j.e. & wouters, b. 2013: evidence and analysis of 2012 greenland records from spaceborne observations, a regional climate model and reanalysis data. th e cryosphere 7, 615–630. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: rsf@geus.dk geological survey of denmark and greenland bulletin 41, 2018, 47-50 47 while multispectral images have been in regular use since the 1970s, the widespread use of hyperspectral images is a relatively recent trend. this technology comprises remote measurement of specific chemical and physical properties of surface materials through imaging spectroscopy. regional geological mapping and mineral exploration are among the main applications that may benefit from hyperspectral technology. minerals and rocks exhibit diagnostic spectral features throughout the electromagnetic spectrum that allow their chemical composition and relative abundance to be mapped. most studies using hyperspectral data for geological applications have concerned areas with arid to semi-arid climates, and using airborne data collection. other studies have investigated terrestrial outcrop sensing and integration with laser scanning 3d models in ranges of up to a few hundred metres, whereas less attention has been paid to ground-based imaging of more distant targets such as mountain ridges, cliffs or the walls of large pits. here we investigate the potential of using such data in well-exposed arctic regions with steep topography as part of regional geological mapping field campaigns, and to test how airborne hyperspectral data can be combined with similar data collected on the ground or from moving platforms such as a small ship. the region between the fjords ikertoq and kangerlussuaq (søndre strømfjord) in west greenland was selected for a field study in the summer of 2016. this region is located in the southern part of the palaeoproterozoic nagssugtoqidian orogen and consists of high-grade metamorphic orthoand paragneisses and metabasic rocks (see below). a regional airborne hyperspectral data set (i.e. hymap) was acquired here in 2002 (tukiainen & thorning 2005), comprising 54 flight lines covering an area of c. 7500 km2; 19 of these flight lines were selected for the present study (fig. 1). the target areas visited in the field were selected on the basis of preliminary interpretations of hymap scenes and geology (korstgård 1979). two different sensors were utilised to acquire the new hyperspectral data, predominantly a specim aisafenix hyperspectral scanner due to its wide spectral range covering the visible to near infrared and shortwave infrared parts of the electromagnetic spectrum. a rikola hyperspectral imager constituted a secondary imaging system. it is much smaller and lighter than the fenix scanner, but is spectrally limited to the visible near infrared range. the results obtained from combining the airborne hyperspectral data and the rikola mineral mapping by hyperspectral remote sensing in west greenland using airborne, ship-based and terrestrial platforms sara salehi and simon mose thaarup ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ;; ; ; ; ; ; ; ; ; ;; ; ; ; ; ; ; ; ; ; sukkertoppen iskappe ;quaternary cover sisimiut intrusive suite amphibolite leucogabbro and anorthosite granodiorite orthogneiss, mainly granodioritic to tonalitic orthogneiss, granulite facies mica schist p a r a u t o c h t h o n o u s a l l o c h t h o n o u s a u t o c h t h o n o u s p a r a u t o c h t h o n o u s a l l o c h t h o n o u s a u t o c h t h o n o u s scan 1 scan 2 scan 3 scan 4scan 5 scan 6 scan 7 scan 8 67 °n 52°w 53°w 66 °3 0' n s ø n d r e s t r ø m f j o r d k a n g e r l u a r s u k ikertooq m ali gia q scan 9 scan 1 scan 2 scan 3 scan 4scan 5 scan 6 scan 7 scan 8 scan 9 qeqerta lik 10 km fig. 1. geological map across the southern nagssugtoqidian orogen simplified from garde & marker (2010) and locations of scanned areas mentioned in the text. red frame: coverage of airborne hyperspectral data selected for this study. from south to north this region comprises (k. sørensen, personal communication 2018): an autochthonous zone in which deformation regularly increases towards the parautochthonous zone expressed in clockwise rotation of kangâmiut basic dykes, a parautochthonous zone in which nagssugtoqidian deformation and metamorphism are highly heterogeneous, and an allochthonous zone where juvenile metasedimentary rocks are interfolded and thrust-stacked with archaean gneisses containing amphibolitic metadykes. © 2018 geus. geological survey of denmark and greenland bulletin 41, 47–50. open access: www.geus.dk/bulletin http://www.geus.dk/bulletin 4848 instrument are presented in salehi (2018), this volume. in addition, representative samples of the main rock types were collected for subsequent laboratory analysis. a parallel study was integrated with geological and 3d photogrammetric mapping in karrat region farther north in west greenland (rosa et al. 2017; fig. 1). analysis of surface mineralogy using subhorizontal hyperspectral data collection the extreme influence of atmospheric effects and topography-induced illumination differences in long-range groundbased spectra data cannot be corrected by means of correction tools commonly used for nadir satellite or airborne data. an adapted workflow is presented in lorenz et al. (2018) to overcome the challenges of long-range outcrop sensing, including atmospheric and topographic corrections. minimum noise fraction transformation (boardman 1993) and spectral angle mapper classification (kruse et al. 1993) were applied to test the applicability of the data for mapping the main rockforming minerals. the former is important for dimensionality reduction and filtering of noise from hyperspectral data. this method can extract spatially coherent information and show the variations between bands in the hyperspectral data. the spectral angle mapper classification permits rapid mapping of the similarity between image and reference spectra. long-range terrestrial hyperspectral scanning in the southern nagssugtoqidian orogen in the summer of 2016, new hyperspectral datasets from relatively distant targets were acquired in the allochthonous (fig. 1, scans 1–6) and autochthonous (fig. 1, scan 8) zones of the fig. 2. a, c: true colour fenix hyperspectral scenes from the maligiaq area. see fig. 1 for scene locations. b, d: minimum noise fraction false colour image: red: band 2. green: band 7. blue: band 8, overlain on grayscale hyperspectral images. facies transition metasedimentary rocks metasedimentary rocks facies transition amphibolite faciesgranulite facies a b c d n s n s 49 nagssugtoqidian orogen. metasedimentary rocks in the allochthonous zone are tectonically interleaved with quartzofeldspathic gneisses between the southern shore of qeqertalik fjord and the northern border of the ikertoq shear zone (figs 1, 2). parts of this zone are well exposed along the eastern shore of maligiaq and were scanned with the fenix instrument from a distance of c. 2–3 km (see fig. 1 and below). the metasedimentary rocks are predominantly biotite-garnet paragneisses but also comprise aluminous schists and graphitic– sulfidic varieties with up to a few percent graphite and iron sulfides, and they may include both archaean and palaeoproterozoic components (see legend to the regional map of garde & marker 2010). the scanned area also transects an amphibolite–granulite facies boundary, which follows one of the shear zone branches (fig. 2). such facies transitions may be difficult to precisely identify in the field, because the occurrence of the granulite facies index mineral hypersthene depends on bulk rock composition in addition to p, t and xh2o conditions; hypersthene first occurs in mafic rocks and will be absent from leucocratic metasedimentary and magmatic rocks at similar metamorphic grade. besides, granulite facies rocks can be retrogressed along younger shear zones. the result generated from the fenix spectral data (fig. 2) highlights the metasedimentary screens. the graphiteand sulphide-bearing schists mapped using such data are shown with purple colour in fig. 2b and pink in fig. 2d. the general colour differences are caused by changes in illumination conditions between the acquisition of the two datasets. while graphitic and sulfidic lithologies are readily detected with the minimum noise fraction method used here, the metamorphic transition from amphibolite facies to granulite facies (korstgård 1979) is not distinguishable; this might be due to absence of diagnostic, hypersthene-bearing lithologies. the boundary might be detected e.g. matched filtering (harsanyi & chang 1994) and mixture-tuned matched filtering (boardman 1998), where changes in mineral chemical composition and information on relative abundances of the lithologies are enhanced. integration of ship-based hyperspectral scanning and 3d photogrammetry mobile mapping of steep coastal cliffs using ship-based hyperspectral scanning was tested – for the first time – in the outer part of the kangerlussuaq area (figs 3, 4) and in the related study in the karrat region mentioned in the introduction. the data processing related to such data acquisition has been fully discussed in salehi et al. (2018). our analysis in the former area reveals kangâmiut dykes intruding intensely deformed archaean gneisses, besides other metabasic bodies (fig. 3). the gneissic foliation is clearly identifiable, and several ductile and brittle structures can be observed in the dykes and other metabasic rocks, where amphiboles are the dominant minerals and mica and pyroxene are present with secondary abundances. the spectral mapping products were integrated with 3d photogrammetric data to create accurate, large-scale outcrop models, which are well suited for quantitative purposes in geological analysis or in preparation for field operations (fig. 4). for this, a dense and accurate dataset of topographic points (also referred to as a point cloud) can be generated from stereo images describing the terrain surface in three dimensions. the implemented algorithms work reliably even for complex geometries, and with high accuracy. slightly distorted data, such as images over a low-relief landscape, can be fig. 3. scan 9 in fig. 1. a: 2d pseudo-orthophoto. b: the resulting spectral angle mapper classification image from the søndre strømfjord region. see salehi et al. (2018) for a detailed description of how to match hyperspectral products to pseudo-orthophotos. spectral end-members used in the classification are shown in the plot. amphibole amphibole + pyroxene host rock (gneiss) mica-rich gneiss amphibole-rich gneiss masked pixels trace of foliation scree archaean gneiss kangâmiut dykes ~ ~~ ~ ~ ~ ~ wavelength (nm) 500 1000 1500 2000 re fle ct an ce (o ffs et fo r c lar ity ) amphibole amphibole + pyroxene host rock (gneiss) amphibole-rich gneiss mica-rich gneiss 5050 treated quickly using homographic or polynomial transformations, and even data with high local distortions caused by the underlying topography can be processed. concluding remarks the workflow presented here for the acquisition of spectral data from moving platforms and long-range, ground-based hyperspectral scanning opens up a range of new possibilities in the application of hyperspectral imagery by significantly enlarging the scale of measurements. the proposed automatic approach to combine spectral and point cloud data is a fast alternative to manual approaches and has high potential for field geologists who wish to establish accurate outcrop models in areas of difficult access that can be brought to life and visualised in 3d surface models. acknowledgments kai sørensen is thanked for constructive discussions on geology. the helmholtz institute freiberg is thanked for the use of the specim aisafenix hyperspectral scanner and the rikola hyperspectral imager. references boardman, j. 1993: automated spectral unmixing of aviris data using convex geometry concepts: in: summaries of the fourth annual jpl airborne geoscience workshop october 25–29, 1993. pasadena, ca: summaries, jpl publication 93-26. vol. 2. boardman, j.w. 1998: leveraging the high dimensionality of aviris data for improved sub-pixel target unmixing and rejection of false positives: mixture tuned matched filtering. in: summaries of the seventh annual jpl airborne geoscience workshop, 97(1), 55–56. garde, a.a. & marker, m. 2010: geological map of greenland, 1:500 000, søndre strømfjord–nuussuaq, sheet 1. second edition. copenhagen: geological survey of denmark and greenland. harsanyi, j.c. & chang, c.i. 1994: hyperspectral image classification and dimensionality reduction: an orthogonal subspace projection approach. ieee transactions on geoscience and remote sensing 32, 779–785. korstgård, j.a. 1979: nagssugtoqidian geology. rapport grønlands geologiske undersøgelse 89, 63–75. kruse, f.a., lefkoff, a., boardman, j., heidebrecht, k., shapiro, a., barloon, p. & goetz, a. 1993: the spectral image processing system (sips) – interactive visualization and analysis of imaging spectrometer data. remote sensing of environment 44, 145–163. lorenz, s., salehi, s., kirsch, m., zimmermann, r., unger, g., sørensen, e.v. & gloaguen, r. 2018: radiometric correction and 3d integration of long-range ground-based hyperspectral imagery for mineral exploration of vertical outcrops. remote sensing 10, 23 pp., http://dx.doi. org/10.3390/rs10020176 rosa, d. et al. 2017: architecture and mineral potential of the palaeoproterozoic karrat group, west greenland. results of the 2016 season. danmarks og grønlands geologiske undersøgelse rapport 2017/114. salehi, s. 2018: hyperspectral analysis of lithologies in the arctic in the presence of abundant lichens. geological survey of denmark and greenland bulletin 41, 51–55 (this volume). salehi, s., lorenz, s., sørensen, e.v., zimmermann, r., fensholt, r., heincke, b.h., kirsch, m. & gloaguen, r. 2018: integration of vessel-based hyperspectral scanning and 3d-photogrammetry for mobile mapping of steep coastal cliffs in the arctic. remote sensing 10, 26 pp, http://dx.doi.org/10.3390/rs10020175 tukiainen, t. & thorning, l. 2005: detection of kimberlitic rocks in west greenland using airborne hyperspectral data: the hypergreen 2002 project. geological survey of denmark and greenland bulletin 7, 69–72. h yp er sp ec tr al d at a c ub e rg b im ag es 1. data collection 2. spectral mapping of lithologies 3. integration with 2d outcrop model 4. projection on 3d topographic model x y z x z 1500 m 10 00 m fig. 4. general workflow for 3d-integration of hyperspectral data cubes and point clouds. authors’ address s.s. & s.m.t., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ssal@geus.dk. http://dx.doi.org/10.3390/rs10020176 http://dx.doi.org/10.3390/rs10020176 http://dx.doi.org/10.3390/rs10020175 mailto:ssal@geus.dk geological survey of denmark and greenland bulletin 1, 931-948 931 shelf-edge delta and slope deposition in the upper callovian – middle oxfordian olympen formation, east greenland michael larsen and finn surlyk the upper bajocian – upper volgian succession of the jameson land basin in east greenland forms an overall transgressive–regressive cycle. the upper callovian – middle oxfordian olympen formation represents the first regressive deposits after maximum flooding in the middle to early late callovian. the formation was deposited during two southwards progradational phases separated by a major drowning event in the early oxfordian. the first phase was marked by incoming of massive slope and base-of-slope sand (athene member), but the delta front and top did not reach the area of present-day exposure. the second phase was initiated by deposition of a thick mud succession (hades member) indicating that the delta had shifted far to the north during the drowning event. southwards progradation of the delta was heralded by gully erosion and the deposition of lenticular bodies of massive slope sand; on this occasion, mediumand largescale cross-bedded sand of the delta front and top (zeus member) reached the area. the boundary between middle–upper callovian mudstones in the upper part of the underlying fossilbjerget formation and the upper callovian athene member sandstones formed at the turn-around point between sea-level rise and fall. the athene member sandstones are interpreted as an undifferentiated falling stage – lowstand systems tract and span a sequence boundary. the top of the athene member is the basinal correlative of the transgressive surface. the basal few metres of the overlying hades member mudstones represent the transgressive systems tract and a level with organic-rich mudstones is interpreted to represent the maximum flooding zone. the remainder of the hades member and the slope sandstones are assigned to the highstand systems tract. the succeeding cross-bedded delta front sandstones of the zeus member are placed in the falling stage systems tract and their sharp base is interpreted as a marine regressive surface of erosion. comparison of this history with published sea-level curves suggests that the short term changes may be eustatic in origin including the middle callovian maximum flooding (k. jason – lower p. athleta chronozones), late callovian regression (p. athleta – q. lamberti chronozones), latest callovian – early oxfordian flooding (q. mariae – c. cordatum chronozones) and late early – middle oxfordian regression (c. densiplicatum chronozone). keywords: east greenland, jameson land basin, middle–upper jurassic, sedimentology, lithostratigraphy, sequence stratigraphy, shelf-edge delta, slope gullies, massive sandstones, sediment gravity flow, sea-level curve m.l., geological survey of denmark and greenland, geocenter copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: mil@geus.dk f.s., geological institute, university of copenhagen, geocenter copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. geological survey of denmark and greenland bulletin 1, 931–948 (2003) © geus, 2003 shelf-edge deltas and their associated slope/base-of-slope deposits are becoming increasingly well-known, especially from the passive margins of the gulf of mexico and the western mediterranean (suter & berryhill 1985; suter et al. 1987; tesson et al. 1990, 1993; sydow & roberts 1994; gensous & tesson 1996; henriksen & weimer 1996). in this study we describe upper callovian – middle oxfordian sandy shelf-edge deltas and slope gravity flow deposits formed during the early stages of jurassic rifting in the east greenland basin. in the jameson land basin, rifting was initiated in the early bajocian, reached a climax in the volgian and waned in the earliest cretaceous. the resulting marine middle – lower upper jurassic rift succession forms an overall transgressive–regressive cycle with a duration of c. 30 ma. the basin axis was oriented north– south, deepening southwards. middle jurassic sediment influx was from the north and transport was mainly axial towards the south resulting in a marked north–south grain-size gradient with a thick succession of shallow marine sandstones in the northern part of the basin passing into thinner offshore mudstones towards the south. the aim of the study is to describe the development of the forestepping late callovian – middle oxfordian shelfedge delta and slope system constituting the olympen formation of the jameson land basin (surlyk et al. 1973; surlyk 2003, this volume, fig. 5). it marks the initiation of the stepwise late jurassic regression after middle to early late callovian maximum flooding of the basin. 932 3 km 71°27'n 23°40'w 23°20´ ice quaternary olympen formation olympen formation fossilbjerget formation lower and middle jurassic undifferentiated olympen 71°24'n sections 1–9 25 km 22°w24°w 71°n jameson land olympen antarctic havn parnas fossilbjerget mikael bjerg 23°30'w parnas o lym pe lve n greenland fig. 1. map showing the distribution of the upper callovian – middle oxfordian olympen formation in jameson land and place names mentioned in the text. geological setting the onshore part of the late paleozoic – mesozoic rift basin of east greenland is about 600 km long and about 200 km wide at the southern end where the jameson land basin is located (fig. 1). the basin was uplifted in tertiary times and the mesozoic succession is excellently exposed. in east greenland, the early jurassic period was tectonically quiescent and deposition was restricted to the jameson land basin, whilst the northern parts of the rift basin were emergent. marine communication between the boreal sea and the tethys ocean was effectively obstructed at the end of the early jurassic by uplift of the north sea dome (surlyk et al. 1973; underhill & partington 1994). distinct faunal provinces rapidly developed and a separate ammonite-based biostratigraphic scheme is used for the bajocian – lower callovian interval in the boreal realm (callomon 1993; 2003, this volume). a combination of bathonian–callovian eustatic sealevel rise, domal deflation and erosion, and onset of the important middle–late jurassic rift phase led to gradual resumption of the north–south marine connection from mid-callovian times. although there was still a marked faunal provincialism in the oxfordian–kimmeridgian, the degree of faunal overlap allows good correlation between east greenland and europe. maximum flooding of the jameson land basin took place in the middle to early late callovian when the middle jurassic sand-dominated system was completely drowned and offshore muds were deposited throughout the basin. rifting increased in the oxfordian– kimmeridgian and reached a climax in the volgian, contemporaneous with the onset of a late jurassic eustatic sea-level fall. this led to a punctuated late jurassic regression and by late volgian times, deltaic deposits had prograded to the southernmost end of jameson land. during the early part of the rift phase, in bajocian– bathonian times, the sea floor was essentially flat and differentiation into shelf, slope and basin was not developed. however, continued southwards axial sediment transport resulted in higher sedimentation rates and preferential sand deposition in the northern part of the basin, whereas the southern part mainly received mud. this led to the gradual development of a northern shallow marine shelf passing southwards into an east–west striking ramp with an incipient slope grading into a deeper-water basin towards the south. two important regressive phases separated by a major drowning event took place in the late callovian and middle oxfordian, respectively, and southwards axial progradation of shallow marine and deltaic sands reached central jameson land. a marked slope was formed in front of the two forestepping successions, and well-developed slope and shelf-edge delta deposits are preserved on several mountain tops in central jameson land (fig. 1). the slope facies are composed of dark grey mudstones and massive, sediment gravity flow sandstones. similar facies types are known from the slightly younger upper oxfordian – volgian hareelv formation in southern jameson land (surlyk 1987) but this latter unit has experienced major post-depositional modification characterised by syndepositional loading structures and evidence of post-burial liquefaction and intrusion of sands into the adjacent and overlying mud (surlyk & noenygaard 2001). post-middle oxfordian shelf-edge deltas are only preserved in the middle–upper volgian part of the succession in southernmost jameson land, due to present-day erosion levels. the shelf-edge delta and slope deposits of the late callovian and middle oxfordian progradational phases constitute the olympen formation of surlyk et al. (1973; fig. 2). the olympen formation stratigraphy the olympen formation was defined by surlyk et al. (1973) for an upper callovian – middle oxfordian tripartite sandstone–mudstone–sandstone succession forming the youngest pre-quaternary sedimentary unit in central jameson land (fig. 1). the formation is divided into three members following the revised lithostratigraphic scheme for the jurassic of east greenland, provisionally introduced by surlyk (2003, this volume, fig. 5). it consists of: (1) a lower unit of massive, fineto medium-grained sandstones intercalated with subordinate, laminated, dark silty mudstones and finegrained laminated sandstones termed the athene member, (2) a middle unit, termed the hades member, of dark silty mudstones which passes upwards into sandy mudstones intercalated with lenticular bodies of massive sandstones and (3) an upper unit of massive or large-scale cross-bedded, mediumto coarse-grained sandstones with subordinate intercalations of silty mudstones termed the zeus member (figs 2, 3; surlyk 2003, this volume, fig. 5). the complete thickness is not known as the formation forms the top of the succession in central jameson 933 934 mud f m c gr 80 70 60 50 40 30 20 10 0 poor exposure poor exposure q. mariae chronozone (early oxfordian) lower–middle p. athleta chronozone (late callovian) p. athleta chronozone (late callovian) k. jason chronozone (middle callovian) c. densiplicatum chronozone (middle oxfordian) ba si n flo or o ly m pe n fo rm at io n fo ss ilb je rg et fo rm at io n sl op e ap ro n a th en e m em be r h ad es m em be r z eu s m em be r ba si n flo or /s lo pe ba si n flo or /s lo pe sl op e gu lly de lta t op de lta t op de lta fr on t lo w er c yc le u pp er c yc le c. 900 m massive sandstone planar cross-bedded sandstone trough cross-bedded sandstone laminated silty mudstone parallel-laminated sandstone carbonate concretion pebbles carbonaceous material ripple formset ripple cross-lamination climbing ripple lamination wave ripple lamination antidunes load structures ripple lamination planar cross-bedding gully axis ammonite belemnite bivalve leaf imprints helminthoida isp. planolites isp. diplocraterion habichi skolithos isp. piscichnus ?kulindrichnus mm 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 fossils facies directional features trace fossils sand mud f m c gr sand 3 7 fig. 2. vertical sections including the type section (left) of the olympen formation at the head of the river olympelven. the olympen formation overlies mudstones of the fossilbjerget formation. note the tripartite lithostratigraphic subdivision of the olympen formation into lower and upper sandstone-dominated units of the athene and zeus members, respectively, separated by the middle mudstone-dominated hades member. the upper boundary is not exposed as the formation forms the summit of the mountains in central jameson land. 935 land. the type section (figs 2–4, section 3) at the top of the olympen mountain measures 150 m, but the thickness in this area may reach 250 m, the upper part being poorly exposed below scree, till and ice. at the parnas, fossilbjerget and mikael bjerg mountains, the formation is 300 m, 250 m and 200 m thick, respectively; further south, in the hurry inlet region, it wedges out completely or is only represented by a few metres of mudstones (surlyk 1991). the olympen formation conformably overlies fossiliferous middle and lower upper callovian mudstones of the fossilbjerget formation of surlyk (2003, this volume, fig. 5; equivalent to the fossilbjerget member of surlyk et al. 1973). the base of the olympen formation is defined by the base of the first thick massive sandstone in the succession (fig. 2). the age of the lower boundary of the formation generally becomes younger from north to south. at the type section, ammonites of the middle callovian k. jason chronozone occur 20 m below the base of the olympen formation and the lower upper callovian p. athleta chronozone is represented 13 m below the base (fig. 2). ammonites from the lower–middle p. athleta chronozone (kosmoceras (zugokosmokeras) cf. proniae teisseyre) occur 3 m below the base, close to the type section, and have also been found 15 m above the base north of olympen (figs 1, 2; birkelund et al. 1971). at mikael bjerg, the upper part of the fossilbjerget formation yields ammonites of the lower p. athleta chronozone (longaeviceras keyserlingi sokolov) in contrast to the northernmost outcrops at antarctic havn where the same ammonite fauna occurs in the lowermost member (athene member) of the olympen formation (callomon 1993). the middle mudstone-dominated hades member has yielded ammonites of the lower oxfordian q. mariae chronozone, c. scarburgense subzone (surlyk et al. 1973; callomon 1993). scattered finds of ammonites from the upper sandstone-dominated zeus member indicate the presence of the middle oxfordian c. densiplicatum chronozone, c. vertebrale subzone (birkelund et al. 1971; callomon 1993). the base of the olympen formation is located just above a regional drowning surface at the turn-around point between the backstepping upper bajocian – upper callovian pelion–fossilbjerget formation couplet and the forestepping upper callovian – middle oxfordian olympen formation (surlyk 1991; 2003, this volume). the olympen formation records two marked progradational phases, one in the late callovian (p. athleta chronozone) and one in the early middle oxfordian (c. densiplicatum chronozone), separated by a drowning event in the early oxfordian (q. mariae chronozone). sedimentary facies recent field work has demonstrated that the upper bajocian – upper volgian package of the jameson land basin forms an overall transgressive–regressive cycle and that massive base-of-slope sandstones are developed in front of stacked, forestepping shelf-edge deltas from the late callovian and through the rest of the jurassic period (surlyk 2003, this volume). the facies of the olympen formation clearly fall within this spectrum of shelf-edge delta and slope deposits and are described below. 1 2 fossilbjerget fm olympen fm athene member olympen fm hades member olympen fm zeus member 3 4 5 6 7 8 9 nesw 50 m fig. 3. photomosaic of the type locality at olympelven on the south side of the glacier-covered mountain olympen. the positions of the measured vertical sections (1–9) are shown (see also fig. 4). 936 mud f m c gr mud f m c gr mud f m c gr mud f m c gr 1 2 3 4 10 20% n = 43 v = 191° 100 m 150 m300 m 500 m south delta top, zeus member planar cross-bedding t st h st hst fs st fs st –l st poor exposures sand sand sand sand fig. 4. north–south correlation panel of the olympen formation at the type locality. sheet-like beds of massive sandstone deposited from sediment gravity flows dominate the athene member. the hades member consists of silty mudstones with lenticular bodies of gravity flow sandstones (sections 7–9) and is overlain by cross-bedded, coarse-grained sandstones of the zeus member. the succession represents two cycles of slope and shelf-edge delta progradation. for legend, see fig. 2; n = number of measurements, -v = vector mean. 937 mud f m c gr mud f m c gr mud f m c gr mud f m c gr mud f m c gr 5 6 7 8 9 1020% n = 26 v = 151° 50 m 400 m 500 m 50 m north turbidites, athene member cross-lamination 10 m rse mfz tlst mfz sand sand sand sand sand hst highstand systems tract fsst falling stage systems tract lst lowstand systems tract tst transgressive systems tract mfz maximum flooding zone tlst top of lowstand surface rse regressive surface of erosion 938 facies 1. laminated mudstones the facies consists of dark brown or dark grey to black, micaceous, silty and sandy mudstones. they are generally well-laminated although the lamination is commonly disrupted by bioturbation. at some levels, the facies is heterolithic with alternating very fine-grained sandstone and mudstone laminae giving the facies a striped appearance (fig. 5a). locally the sandstone laminae pinch and swell with incipient ripple formsets some of which show cross-lamination. helminthoida isp. grazing or crawling traces occur in high densities on bedding planes. the mudstones have a total organic carbon (toc) content of 1–9% and a total sulphur (ts) content of 0.1–1.3%. the highest toc value is shown by a black finely laminated mudstone 2 m above the base of the hades member (fig. 2, section 3, 78 m). carbonate concretions are common and form isolated lenses or lensoid layers. they have nuclei of fossil wood and, in a few cases, ammonites. the mudstones contain rare belemnites, bivalves and ammonites. a marine, low diversity dinoflagellate cyst assemblage was described by fensome (1979). facies 1 is dominated by fine-grained sediment deposited from suspension. intervals of ripple cross-laminated siltstones and fine sandstones, however, indicate traction current action representing low-energy turbidity currents or weak bottom currents. facies 2. cross-laminated sandstones the facies consists of well-sorted, very fineor finegrained current ripple cross-laminated sandstone. set fig. 5. a: heterolithic mudstones and fine-grained sandstones (facies 1) intercalated with thin-bedded, massive sandstones (facies 4). the upper massive sandstone is strongly erosional and probably represents a local cut-and-fill. athene member, section 4. pencil is 14 cm long. b: fine-grained sandstones showing climbing ripple cross-lamination (facies 2). the sandstones are closely associated with laminated mudstones and parallel-laminated sandstones and represent low-density turbidites. scale is 10 cm long. a b 939 thickness normally varies from 0.5–2 cm, but may reach 4 cm. asymmetric ripple-formsets and climbing-ripple cross-lamination occur locally (fig. 5b). ripple asymmetry and orientations of foresets indicate a south to south-easterly palaeocurrent direction (fig. 4). the facies forms sandstone-dominated packages with gradational boundaries to facies 1. the rippled sandstones were deposited from low-density turbidity currents (bouma 1962) or weak bottom currents. the close association with mudstones of facies 1 suggests that the cross-laminated sandstones were deposited in a deep-water environment with only episodic sand deposition. facies 3. parallel-laminated sandstones the facies is composed of well-sorted, very fine-grained or fine-grained sandstones forming beds up to 60 cm thick. these beds show marked lateral changes in thickness from a few centimetres to 60 cm over a few tens of metres, and pinch-and-swell morphology is common. the thicker beds show parallel and low angle hummocky-like stratification associated with soft-sediment deformation structures including overturned folds. antidune lamination showing aggradation on the upslope side of low-angle bedforms and draping of previous topography by parallel-laminated sandstones is also represented. the parallel-laminated sandstones mainly have flat bases, but scouring occurs locally. the erosional topography is filled by laminated sandstones with divergent laminae thickening over the deepest part of the scour depressions. facies 3 is interpreted as having been deposited from low density turbidity currents (tb division; bouma 1962). some of the structures show a superficial resemblance to hummocky cross-stratification, but are interpreted as having been formed in the upper flow regime with plane bed deposition transitional to antidune bedding (skipper 1971; hand et al. 1972; prave & duke 1990). facies 4. massive sandstones the facies consists of white to light grey, micaceous, fineto medium-grained well sorted, massive sandstone beds commonly with a slightly graded, parallel-laminated top rich in disseminated plant material. two subfacies are recognised. subfacies 4a. sheet-like massive sandstones the subfacies includes sheet-like units of massive sandstones, up to 8 m thick, built up of amalgamated beds 0.5–2 m thick, or occasionally comprising single beds. the lower boundaries of the sandstone sheets are sharp, but typically flat and apparently non-erosional (fig. 6). the sandstone beds in general show parallel boundaries, but emplacement folds and local scouring sometimes occur at amalgamation surfaces (fig. 7a; heller & dickinson 1985). the individual sandstone beds are massive with grading in the uppermost few centimetres. the graded tops are rich in disseminated, carbonaceous plant material and commonly show deformation by water-escape and loading (fig. 7b). amalgamated beds may split into separate beds over a distance of a fig. 6. sheet-like units of amalgamated, massive, fine-grained sandstone beds (facies 4a) interbedded with mudstonedominated intervals (facies 1–3) in the athene member. note the sharp but apparently non-erosional lower boundary of the sandstones (arrows). the white sandstone bed in the centre of the photo is 8 m thick. section 1, 15–40 m above base of athene member (fig. 4). 940 2 cm fig. 7. a: amalgamated, massive, finegrained sandstone beds separated by thin carbonaceous levels in the lower part of the athene member (facies 4a). note the truncation of the lower beds in the centre of the photograph (arrows). person for scale. b: top of a massive sandstone bed (facies 4). note the horizon of carbonaceous plant material showing flow structures (arrow) overlain by parallel-laminated fine-grained sandstones. the top part shows a concentration of low-density organic material and the upwards change in structures suggest deposition from the waning phase of a turbidity current. the sharp base of the overlying bed is indicated by a dashed line. section 3. fig. 8. massive sandstone bed (facies 4), 12 m thick, forming the top of the athene member. the massive nature of the sandstone suggests deposition from either a sandy debris flow or a sustained high-density turbidity current. section 5; person (encircled) for scale. a b 941 few hundred metres. pyrite concretions, up to 3 cm in diameter, are common along the lower bed boundaries. in the upper part of the sandstone-dominated athene member, an up to 17 m thick massive sandstone bed is referred to facies 4a (fig. 8). the base of the bed shows evidence of loading, but only minor erosion. the massive sandstones are interpreted as the deposits of sediment gravity flows, either turbidity currents or sandy debris flows (lowe 1982; surlyk 1987; kneller & branney 1995). the graded tops with abundant carbonaceous plant material, however, suggest deposition from turbidity currents. this implies that the thick, ungraded part of each bed was probably formed by gradual aggradation of sand beneath a sustained steady or quasi-steady current (kneller & branney 1995). the largely planar and parallel lower and upper bed boundaries indicate that the flows were in a largely non-erosive stage and deposition probably occurred in a base-of-slope and basin floor setting (heller & dickinson 1985). the thicker beds commonly consist of amalgamated units, as indicated by horizons of carbonaceous plant material. emplacement folds between amalgamated sandstones suggest that the sands of the upper bed were emplaced upon the still largely unconsolidated substratum of the underlying bed (heller & dickinson 1985). the common occurrence of packets of amalgamated sandstones separated by thick mudstone units suggests that sandstone deposition occurred in discrete pulses separated by longer periods of background sedimentation of fines. fig. 9. gully fill (facies 4b) at the top of the middle unit (sections 7–9). a: large slab of mudstone lifted by intrusive sand at the base of the gully. the left side of the slab is partly in situ, whereas the right side is lifted and truncated. b: mudstone rip-up clast in the lower part of the gully fill sandstone. note the sharp lower boundary of the sandstone (arrows); hammer for scale. a b 942 subfacies 4b. lenticular massive sandstones the subfacies consists of lenticular sandstone bodies, up to 16 m thick, that fill deep erosional scours or gullies several tens of metres wide in cross-section. the most spectacular example occurs in the upper part of the hades member (fig. 4, section 7). at this locality, a large slab of mudstone is partially detached, and truncated by massive sands at the base of the gully fill (fig. 9a). the sandstone fill is massive, fineto medium-, or locally coarse-grained, and shows normal grading in the lower levels and at the top. floating well-rounded, oblate quartzite pebbles occur in the lower levels of one thick gully fill (fig. 2, section 7, 30 m). mudstone clasts ranging from a few millimetres to several metres in largest dimension occur scattered or in distinct horizons (fig. 9b). the larger clasts are found relatively close to the base of the sandstone beds. the mudstone clasts are commonly undeformed except for torn-up ends (fig. 10a), but sheared clasts occur along the lower margins of some beds. the axis of the gully exposed in section 7 strikes nw–se (137°). the massive sandstones grade upwards into parallel-laminated sandstones from a few centimetres to several metres thick, capped by cross-laminated sandstones (fig. 10b). the massive sandstone bodies were emplaced by sediment gravity flows (lowe 1982; kneller & branney 1995). the basal grading and the upwards change into parallel-laminated sandstones suggest deposition from sustained high-density turbidity currents (kneller & branney 1995) although a sandy debris flow origin is also possible. horizons of mudstone intraclasts may mark the forfig. 10. a: zones with aligned mudstone clasts in the lower part of the gully fill indicate that deposition occurred in pulses or as separate flows that followed shortly after one another. b: several metres thick unit of parallel-laminated fine-grained sandstones forming the top part of the gully fill, section 7. hammer (centre right) for scale. a b 943 mer position of the rising depositional flow boundary or may, in some cases, show the presence of thinner depositional units separated by subhorizontal amalgamation surfaces. the shape of the clasts shows that the adjacent mudstones were cohesive or partly consolidated at the time of gully erosion and sandstone deposition. facies 5.trough cross-bedded sandstones the facies consists of trough cross-bedded, mediumto coarse-grained sandstones with sets 10–15 cm thick, forming cosets up to 1.5 m thick. trough axes indicate palaeocurrent directions towards the south-east. bioturbation is common and includes diplocraterion habichi and skolithos isp. the facies is confined to the zeus member (fig. 2). the trough cross-sets are commonly located at the down-current termination of lenticular bodies of large-scale cross-bedded sandstones of facies 6. the facies was deposited by traction currents in the upper part of the lower flow regime. the trace fossil assemblage suggests a shallow marine environment. facies 6. large-scale cross-bedded sandstones the facies consists of coarse-grained sandstones showing large-scale cross-bedding with tangential, locally sigmoidal foresets (fig. 11). the thickness of the crossbeds varies from 0.4–6 m, but is mainly 2–4 m. the crosssets are composed of graded avalanche foresets that reach maximum dips of 26°. they pass downwards into carbonaceous, strongly bioturbated toesets. double mud drapes are found locally in the sandstones, but no systematic changes were recorded in foreset thicknesses. the cross-sets are commonly truncated at the top and are overlain by trough cross-bedded sandstones. in the upper part of the zeus member, wave-rippled topsets occur. dip azimuths of foresets indicate unimodal palaeocurrent directions towards the south with a vector mean of 191° (fig. 4). bioturbation is concentrated along set boundaries and shows high density and diversity. the trace fossil assemblage includes horizontal burrows of planolites isp., taenidium serpentinum, gyrochorte comosa, vertical burrows of diplocraterion habichi, skolithos isp. and the resting trace piscichnus ?kulindrichnus. the facies is only present in the zeus member (fig. 2). the large-scale cross-sets may pass downcurrent into trough cross-bedded sandstones of facies 5. the cross-beds were formed by southwards progradation of subtidal sand bodies, the top of which were periodically exposed to wave reworking indicating a water depth around wave base. comparison with similar facies in the volgian raukelv formation of southern jameson land (surlyk & noe-nygaard 1991) suggests that the cross-beds may form intrasets in large-scale compound foreset beds. facies successions the facies of the olympen formation form an overall progradational megacycle that is made up of two shallowing-upwards cycles: (1) the uppermost fossilbjerget formation together with the athene member and (2) fig. 11. cross-bedded, coarse-grained sandstones (facies 6) of the zeus member. individual sets are up to 3 m thick (locally up to 6 m), and are separated by heterolithic mudstones and fine-grained sandstones. foresets are commonly tangential with strongly bioturbated toesets and truncated topsets. section 3. 2 m the hades and zeus members of the olympen formation (fig. 2). they are described in turn and interpreted within the framework of the delta-fed turbidite ramp model of heller & dickinson (1985) and surlyk (1987). lower cycle the lower shallowing-upwards cycle is formed by the silty mudstones of the uppermost part of the fossilbjerget formation and the sharp-based, mainly massive sandstones (facies 4) intercalated with mudstones (facies 1) and fine-grained sandstones (facies 2 and 3) forming the athene member (fig. 2). the first sandstone bed is 4 m thick and abruptly overlies a uniform succession of silty mudstones, several tens of metres thick. above this level, massive sandstones of facies 4 dominate volumetrically. they are commonly amalgamated and may form units up to 17 m thick, whereas single beds may reach 8 m in thickness. the laterally persistent sheetlike beds of massive sandstone (facies 4a) dominate the lower part of the cycle, whereas the upper part of the cycle is characterised by thick, lenticular sandstones (facies 4b). a few beds of ripple cross-laminated sandstones (facies 2) and parallel-laminated sandstones (facies 3) are intercalated with mudstone units in the middle part of the cycle. the mudstones change from non-bioturbated to bioturbated c. 20 m above the base of the first sandstone bed. a marked change in facies occurs at the transition to the overlying mudstone-dominated association of the hades member and the top of the uppermost massive sandstone bed has been reworked into low-angle inclined sets. at locality 4 (fig. 4), an isolated cross-bedded coarse-grained sandstone bed is intercalated in mudstones, c. 2 m above the base of the upper cycle. the sandstones at the top of the lower cycle are intensively bioturbated and show a diverse trace fossil assemblage compared to the underlying sandstones. upper cycle the base of the upper cycle is marked by a sharp boundary between sandstones at the top of the athene member and dark silty mudstones of the hades member (facies 1; figs 2, 4). the mudstone-dominated lower portion is about 50 m thick and shows a slight coarsening-upwards trend in the upper 15 m. the most finegrained mudstones occur at 1.5 m and 6 m above the base of the cycle and are black, organic-rich (9% toc) and finely laminated. with these exceptions, the micarich mudstones are typically silty, locally sandy and intensely bioturbated. at locality 7 (fig. 4), the mudstones are succeeded by about 16 m of massive, fineto mediumgrained sandstones (facies 4b), which fill a deep erosional scour or gully cut into the mudstones (figs 4, 9). elsewhere, the mudstone succession is sharply overlain by a coarse-grained sandstone succession composed of stacked sets of large-scale cross-bedded sandstones (facies 6), trough cross-bedded sandstones (facies 5) and intercalated mudstones (facies 1; figs 2, 4). channelling and scouring are common. the coarse-grained sandstones are strongly bioturbated and show a diverse trace fossil assemblage. in the uppermost part of section 3 (fig. 4), wave-ripple formsets are common on the upper bedding planes of large-scale cross-sets. the uppermost part of the upper cycle is poorly exposed. isolated outcrops suggest, however, that the coarsegrained cross-bedded sandstone facies continues to the top of the formation. depositional model for the olympen formation the lower cycle represents late callovian progradation of a proximal basin and base-of-slope setting. the dominance of laterally continuous, massive, non-erosional gravity flow sandstones (facies 4a), and the lack of obvious vertical organisation of the sandstones suggest that they represent a proximal ramp facies of a deltafed turbidite system (heller & dickinson 1985; surlyk 1987). the shallow marine part of the delta front did not reach the present outcrop area and the interpretation is thus based on comparison with the upper cycle and with shelf-edge deltas and redeposited mass-flow sandstones of the volgian raukelv formation (surlyk & noe-nygaard 1995). the abrupt incoming of sandstones in the basin plain facies contrasts with the model of heller & dickinson (1985), which predicts a gradual increase in bed thickness and grain size from the basin plain – distal ramp facies to the proximal ramp. the concentration of pyrite concretions in the lower part of the cycle and the upwards change towards more bioturbated sediments suggests a change from a poorly oxygenated deeper-water setting into a more well-aerated, relatively shallow-water environment. a major drowning event occurred in latest callovian – early oxfordian times before the inferred shelf-edge delta of the lower cycle had reached the area of present-day outcrop. this was followed by renewed progra944 945 dation in the early–middle oxfordian. the relatively great thickness of the lower oxfordian basinal mudstones, which form the lower 50 m of the upper progradational cycle, suggests that the drowning event had translated the shelf edge far northwards. at olympen, shelf-edge delta progradation was heralded by the incoming of thick amalgamated slope sandstones with a strongly scoured base. they were probably deposited in a slope gully of the same type as described from the upper oxfordian part of the hareelv formation further south in jameson land (surlyk 1987; surlyk & noenygaard 2001). finally the shelf-edge delta reached the outcrop area as marked by the incoming of coarsegrained high-angle cross-beds. such beds are characteristic of lowstand deltas that prograded to the edge of the shelf in the middle jurassic (pelion formation; engkilde & surlyk 2003, this volume) and especially in the volgian (raukelv formation; surlyk & noe-nygaard 1991, 1995; surlyk et al. 1993; surlyk 2003, this volume). sequence stratigraphy a sequence stratigraphic interpretation of the olympen formation is not straightforward due to the relatively deep-water nature of most of the succession. the mudstones at the base of the lower cycle represent late callovian maximum flooding of the whole jameson land basin. the delta top of the lower unit is not preserved in jameson land. the slope and base-of-slope mass-flow sandstones were most likely deposited during late fall, maximum lowstand or early rise. thus, the base of the lower sandstone package may not necessarily represent the time of maximum sea-level lowstand and the correlative surface to the subaerial sequence boundary could be found within the sandstone package (hunt & tucker 1993, 1995). however, it has not been possible to define this deep-water correlative surface in the succession and present-day exposures do not allow tracing of the surface from the basin margin into the slope system. in our view, the sandstone succession is thus best interpreted as an undifferentiated falling stage and lowstand systems tract (fig. 4). the drowning surface at the top of the lower cycle is interpreted to represent the slope correlative of the transgressive surface defining the top of the lowstand systems tract. it probably passes up-dip into a ravinement surface, which extended across the inner shelf to the north. the thin organic-rich mudstone unit in the basal part of the second cycle represents the maximum flooding zone overlying a transgressive systems tract which is only a few metres thick (fig. 4). it is overlain by a thick coarsening-upwards mudstone-dominated highstand systems tract. the incoming of massive sandstones indicates that the delta had prograded to the shelf edge, and massive sands were shed down the slope from the delta front. the sharp base of the overlying cross-bedded delta front sandstones was possibly formed by marine erosion in front of the prograding delta and may represent a marine regressive surface of erosion (posamentier et al. 1992). the cross-bedded delta front sandstones (fig. 4) are thus tentatively placed in the falling stage systems tract (hunt & tucker 1993, 1995; nummedal et al. 1993). comparison with eustatic sea-level curves since the seminal work of vail et al. (1977; see miall 1997 for an overview), the validity of global sea-level curves has been extensively debated. however, as data from different regions are added and the dating of observed sea-level changes is refined, portions of the curve may prove to be eustatic. a regional sea-level curve for the jurassic of east greenland was presented by surlyk (1990) and compared with the eustatic curves of haq et al. (1988) and hallam (1988). sea-level changes based on the sequence stratigraphic interpretation of the olympen formation can be tied to ammonite chronozones and allow refinement of the regional sea-level curve for the late middle and early late jurassic of the jameson land basin of east greenland (fig. 12). the late jurassic part of the curve is supplemented with data from the charcot bugt formation in milne land situated at the western margin of the east greenland basin (larsen 1995; larsen et al. 2003, this volume). the sea-level changes recorded in the olympen formation include a sea-level rise in the middle callovian (k. jason chron), a late callovian fall (early–middle p. athleta and possibly q. lamberti chrons), an end callovian – early oxfordian rise (q. mariae chron) and a late early – middle oxfordian fall (c. densiplicatum chron) (fig. 12). these changes can be matched closely with a proposed eustatic sea-level curve for the central part of the russian platform (sahagian et al. 1996). the callovian part of the curve does not match the sealevel curve of haq et al. (1988), whereas the early oxfordian rise followed by middle oxfordian fall is recorded in all three curves (fig. 12). conclusions the upper bajocian – upper volgian succession in the jameson land basin, east greenland forms a long-term transgressive–regressive cycle with maximum flooding in the middle callovian and early kimmeridgian and regressive pulses in the late callovian, and the middle and late oxfordian. the upper callovian – middle oxfordian olympen formation represents the initial regression following middle callovian flooding. the formation consists of a tripartite sandstone–mudstone–sandstone package overlying upper callovian mudstones of the fossilbjerget formation. the top of the fossilbjerget formation and the lower sandstone-dominated athene member of the olympen formation together record late callovian progradation of the shelf-edge delta and slope towards the south. the delta top did not reach the area of present-day exposure, however, and the sandstone unit consists solely of massive gravity flow sandstones deposited on the slope and base-of-slope. progradation was terminated by a major early oxfordian drowning event. renewed progradation took place in the early–middle oxfordian heralded by slight coarsening-upwards of the hades member mudstones and the formation of erosional slope gullies filled with massive gravity flow sandstones, tens of metres thick. they are directly overlain by delta front and coarse-grained, cross-bedded delta top sandstones of the zeus member indicating that the delta finally reached the area. the athene member of the olympen formation represents an undifferentiated falling stage – lowstand systems tract. the drowning surface at the top of the athene member is interpreted as the top lowstand surface forming the distal correlative of the transgressive surface. a maximum flooding zone is recognised in the basal part of the mudstones of the hades member, followed by a thick, coarsening-upwards mudstone succession with lenticular bodies of massive sandstones representing the highstand systems tract. the sharp base of the overlying cross-bedded delta front sandstones of the zeus member is interpreted to represent a marine regressive surface of erosion, and the sandstones are tentatively placed in a falling stage systems tract. comparison of this history of middle callovian maximum flooding (k. jason chron), late callovian fall (early–middle p. athleta and possibly q. lamberti chrons), latest callovian – early oxfordian flooding (q. mariae chron) and late early – middle oxfordian (c. densiplicatum chron) regression with published eustatic sea-level curves suggests that the short-term changes can be matched within the resolution of ammonite zones. 946 geochronology haq et al. 1988 sahagian et al. 1996 this study chronozones g. transversarium c. tenuiserratum c. densiplicatum c. cordatumc. cordatum p. plicatilis q. mariaeq. mariae q. lambertiq. lamberti p. athleta short-term long-term p. athleta e. coronatume. coronatum k. jason s. calloviense k. jason s. calloviense chronozones low high low high low highageperiod submediterranean province boreal–subboreal province o xf or di an la te ju ra ss ic m id dl e m id dl e m id dl e la te ea rl y ea rl y c al lo vi an ammonite dating fig. 12. comparison of sea-level curves for the callovian–oxfordian time interval based on the data in this study (east greenland), haq et al. (1988) (global) and sahagian et al. (1996) (russian platform). note the close correlation between short-term changes in the east greenland and russian platform curves. amplitudes of the curves have been rescaled for illustration purposes. ammonite zonation based on sykes & callomon (1979) and callomon (2003, this volume). 947 acknowledgements m.l. gratefully acknowledges a three-year ph.d. stipend from the carlsberg foundation (91-0683/20, 92-0505/20 and 93-0735/20). the danish natural science research council supported subsequent work. we wish to thank michael engkilde and john h. callomon for sedimentological and stratigraphic discussions and the referees, stephen p. hesselbo, kevin t. pickering and duncan pirrie, for their thorough reviews and suggestions to improve the manuscript. references birkelund, t., håkansson, e. & surlyk, f. 1971: new finds of bathonian, callovian and oxfordian ammonites in northern jameson land, east greenland. bulletin of the geological society of denmark 20, 240–259. bouma, a.h. 1962: sedimentology of some flysch deposits: a graphic approach to facies interpretation, 168 pp. amsterdam: elsevier. callomon, j.h. 1993: the ammonite succession in the middle jurassic of east greenland. bulletin of the geological society of denmark 40, 83–113. callomon, j.h. 2003: the middle jurassic of western and northern europe: its subdivisions, geochronology and correlations. in: ineson, j.r. & surlyk, f. 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(eds): oscar volume. bulletin of the geological society of denmark 48, 169–188. surlyk, f., callomon, j.h., bromley, r.g. & birkelund, t. 1973: stratigraphy of the jurassic – lower cretaceous sediments of jameson land and scoresby land, east greenland. bulletin grønlands geologiske undersøgelse 105, 76 pp. surlyk, f., noe-nygaard, n. & dam, g. 1993: high and low resolution sequence stratigraphy in lithological predictions – examples from the mesozoic around the northern north atlantic. in: parker, j.r. (ed.): petroleum geology of northwest europe: proceedings of the 4th conference, 199–213. london: geological society. suter, j.r. & berryhill, h.l. 1985: late quaternary shelf-margin deltas, northwest gulf of mexico. american association of petroleum geologists bulletin 69, 77–91. suter, j.r., berryhill, h.l. & penland, s. 1987: late quaternary sea level fluctuations and depositional sequences, southwest louisiana continental shelf. in: nummedal, d., pilkey, o.h. & howard, j. 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(eds): siliciclastic sequence stratigraphy: recent developments and applications. american association of petroleum geologists memoir 58, 449–484. vail, p.r., mitchum, r.m., todd, r.g., widmier, j.m., thompson, s., sangree, j.b., bubb, j.n. & hatlelid, w.g. 1977: seismic stratigraphy and global changes of sea level. american association of petroleum geologists memoir 26, 49–212. 948 manuscript received 8 january 1998; revision accepted 23 february 1999. danmarks og grønlands geologiske undersøgelse (geus) geological survey of denmark and greenland øster voldgade 10, dk-1350 copenhagen k denmark geological survey of denmark and greenland bulletin is a new series started in 2003 to replace the two former bulletin series of the survey, viz. geology of greenland survey bulletin and geology of denmark survey bulletin. the twenty-one volumes published since 1997 in those two series are listed below, followed by titles in the new bulletin series. the new series, together with geological survey of denmark and greenland map series, now form the peer-review scientific series of the survey. geology of greenland survey bulletin (discontinued) 173 cambrian shelf stratigraphy of north greenland, 120 pp., 1997. by j.r. ineson & j.s. peel. 250.00 174 the proterozoic thule supergroup, greenland and canada: history, lithostratigraphy and development, 150 pp., 1997. by p.r. dawes. 300.00 175 stratigraphy of the neill klinter group; a lower – lower middle jurassic tidal embayment succession, jameson land, east greenland, 80 pp., 1998. by g. dam & f. surlyk. 250.00 176 review of greenland activities 1996, 112 pp. (18 articles), 1997. edited by a.k. higgins & j.r. ineson. 200.00 177 accretion and evolution of an archaean high-grade grey gneiss – amphibolite complex: the fiskefjord area, southern west greenland, 115 pp., 1997. by a.a. garde. 200.00 178 lithostratigraphy, sedimentary evolution and sequence stratigraphy of the upper proterozoic lyell land group (eleonore bay supergroup) of east and north-east greenland, 60 pp., 1997. by h. tirsgaard & m. sønderholm. 200.00 179 the citronen fjord massive sulphide deposit, peary land, north greenland: discovery, stratigraphy, mineralization and structural setting, 40 pp., 1998. by f.w. van der stijl & g.z. mosher. 200.00 180 review of greenland activities 1997, 176 pp. (26 articles), 1998. edited by a.k. higgins & w.s. watt. 200.00 181 precambrian geology of the disko bugt region, west greenland, 179 pp. (15 articles), 1999. edited by f. kalsbeek. 240.00 182 vertebrate remains from upper silurian – lower devonian beds of hall land, north greenland, 80 pp., 1999. by h. blom. 120.00 183 review of greenland activities 1998, 81 pp. (10 articles), 1999. edited by a.k. higgins & w.s. watt. 200.00 184 collected research papers: palaeontology, geochronology, geochemistry, 62 pp. (6 articles), 1999. 150.00 185 greenland from archaean to quaternary. descriptive text to the geological map of greenland, 1:2 500 000, 93 pp., 2000. by n. henriksen, a.k. higgins, f. kalsbeek & t.c.r. pulvertaft. 225.00 186 review of greenland activities 1999, 105 pp. (13 articles), 2000. edited by p.r. dawes & a.k. higgins. 225.00 187 palynology and deposition in the wandel sea basin, eastern north greenland, 101 pp. (6 articles), 2000. edited by l. stemmerik. 160.00 188 the structure of the cretaceous–palaeogene sedimentary-volcanic area of svartenhuk halvø, central west greenland, 40 pp., 2000. by j. gutzon larsen & t.c.r. pulvertaft. 130.00 189 review of greenland activities 2000, 131 pp. (17 articles), 2001. edited by a.k. higgins & k. secher. 160.00 190 the ilímaussaq alkaline complex, south greenland: status of mineralogical research with new results, 167 pp. (19 articles), 2001. edited by h. sørensen. 160.00 191 review of greenland activities 2001, 161 pp. (20 articles), 2002. edited by a.k. higgins, k. secher & m. sønderholm. 200.00 geology of denmark survey bulletin (discontinued) 36 petroleum potential and depositional environments of middle jurassic coals and non-marine deposits, danish central graben, with special reference to the søgne basin, 78 pp., 1998. by h.i. petersen, j. andsbjerg, j.a. bojesen-koefoed, h.p. nytoft & p. rosenberg. 250.00 37 the selandian (paleocene) mollusc fauna from copenhagen, denmark: the poul harder 1920 collection, 85 pp., 2001. by k.i. schnetler. 150.00 geological survey of denmark and greenland bulletin (new series) 1 the jurassic of denmark and greenland, 948 pp., 2003. 500.00 edited by j.r. ineson & f. surlyk. 2 fish otoliths from the paleocene of denmark, 94 pp., 2003. by w. schwarzhans. forthcoming volumes late quaternary environmental changes recorded in the danish marine molluscan faunas. by k.s. pedersen. the jurassic of north-east greenland. edited by l. stemmerik & s. stouge. review of survey activities, 2003. edited by m. sønderholm & a.k. higgins. prices are in danish kroner exclusive of local taxes, postage and handling geological survey of denmark and greenland bulletin 1, 631-656 631 early and middle jurassic mires of bornholm and the fennoscandian border zone: a comparison of depositional environments and vegetation henrik i. petersen, lars h. nielsen, eva b. koppelhus and henning s. sørensen suitable climatic conditions for peat formation existed during early–middle jurassic times in the fennoscandian border zone. autochthonous peat and allochthonous organic matter were deposited from north jylland, south-east through the kattegat and øresund area, to skåne and bornholm. the increase in coal seam abundance and thickness from north jylland to bornholm indicates that the most favourable peat-forming conditions were present towards the south-east. peat formation and deposition of organic-rich muds in the early jurassic coastal mires were mainly controlled by a continuous rise of relative sea level governed by subsidence and an overall eustatic rise. watertable rise repeatedly outpaced the rate of accumulation of organic matter and terminated peat formation by lacustrine or lagoonal flooding. organic matter accumulated in open-water mires and in continuously waterlogged, anoxic and periodically marine-influenced mires. the latter conditions resulted in huminite-rich coals containing framboidal pyrite. the investigated lower jurassic seams correspond to peat and peaty mud deposits that ranged from 0.5–5.7 m in thickness, but were generally less than 3 m thick. it is estimated that on bornholm, the mires existed on average for c. 1200 years in the hettangian–sinemurian and for c. 2300 years in the late pliensbachian; the early jurassic (hettangian–sinemurian) mires in the øresund area existed for c. 1850 years. aalenian uplift of the ringkøbing–fyn high and major parts of the danish basin caused a significant change in the basin configuration and much reduced subsidence in the fennoscandian border zone during the middle jurassic. this resulted in a more inland position for the middle jurassic mires which on occasion enabled peat accumulation to keep pace with, or temporarily outpace, watertable rise. thus, peat formation was sometimes sustained for relatively long periods, and the mires may have existed for up to 7000 years in the øresund area, and up to 19 000 years on bornholm. the combination of the inland position of the mires, a seasonal climate, and on occasion a peat surface above groundwater level caused temporary oxidation of the peat surfaces and formation of inertinite-rich coals. the spore and pollen assemblages from coal seams and interbedded siliciclastic deposits indicate that the dominant plant groups in both the early and middle jurassic mires were ferns and gymnosperms. however, significant floral differences are evident. in the lower jurassic coals, the palynology testifies to a vegetation rich in cycadophytes and coniferophytes (taxodiaceae family) whereas club mosses were of lesser importance. conversely, in the middle jurassic coals, the palynology indicates an absence of cycadophytes, a minor proportion of coniferophytes (taxodiaceae) and a significant proportion of club mosses. these variations are probably related to adaptation by different plants to varying environmental conditions, in particular of hydrological character. keywords: danish basin, fennoscandian border zone, bornholm, lower–middle jurassic, coal distribution, coal petrography, sedimentology, palynology, mire environments, mire vegetation, peat accumulation rates, sea-level change h.i.p., l.h.n., e.b.k.* & h.s.s.**, geological survey of denmark and greenland, geocenter copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: hip@geus.dk present addresses: *royal tyrrell museum of paleontology, box 7500, drumheller t07 0y0, canada. **danfoss a/s, l7-s38, dk-6430 nordborg, denmark. geological survey of denmark and greenland bulletin 1, 631–656 (2003) © geus, 2003 632 in contrast to the permo-carboniferous and cenozoic, the jurassic is not considered to be a time of major peat formation. however, the jurassic succession of northwest europe, and in particular the middle jurassic, contains significant coal-bearing strata showing that favourable conditions for peat formation were fulfilled at many times. in the danish area, coal-bearing strata are generally restricted to the central graben in the north sea and the fennoscandian border zone. well-developed lower–middle jurassic coals are found from the øresund area to the island of bornholm in the baltic sea indicating climatic and geological conditions suitable for the accumulation and preservation of organic matter (fig. 1). the importance of coal beds in stratigraphic analysis is emphasised by the fact that coal seams may represent a considerable proportion of the time represented by coal-bearing successions. thus, this paper views the peat-forming environments on bornholm and in the fennoscandian border zone in a broad geological perspective and discusses the variability in peat-forming conditions in the area in an attempt to explain the principal compositional differences between the lower and middle jurassic coal seams. in addition, the paper presents palynological data from the coals and adjacent sediments and discusses the palynological evidence of the peat-forming vegetation compared to the evidence from the coal composition. geological setting the fennoscandian border zone is a major tectonic structure that separates the baltic shield from the danish basin and polish trough (fig. 1). the border zone is divided into the stable skagerrak–kattegat platform and the highly block-faulted sorgenfrei–tornquist zone, which has been tectonically active at least since late palaeozoic times (sorgenfrei & buch 1964; norling & bergström 1987; eugeno-s working group 1988; michelsen & nielsen 1991, 1993; mogensen 1994). towards the south-east, the sorgenfrei–tornquist zone continues into the teisseyre–tornquist zone via the rønne graben, west of bornholm in the baltic sea. the rønne graben pull-apart basin was formed by late carboniferous – early permian dextral wrench-faulting (vejbæk 1985; liboriussen et al. 1987). continued triassic jylland børglum fault frederikshavn-3 skagen-2 100 km ringkøbing–fyn high bornholm skåne danish basin fjerritslev fault skagerrak–kattegat platform sorgenfrei–tornquist zone well normal fault erosion limit of triassic–jurassic areas of late cretaceous and early tertiary inversion basement high rø nn e g ra be n skagerrak kattegat 56°n 57°n 55°n east north sea high 5 157 18 haldager fault øresund wells øresund 13 8°e 10°e 12°e12°e 14°e fig. 1. map showing the structural outline of the fennoscandian border zone and the location of bornholm and the øresund, skagen-2 and frederikshavn-3 wells. modified from liboriussen et al. (1987) and eugeno-s working group (1988). 633 – early cretaceous wrench faulting resulted in faultcontrolled subsidence and tilting of fault blocks in the sorgenfrei–tornquist zone. compressional tectonism in the late cretaceous – palaeogene led to inversion of fault blocks in the sorgenfrei–tornquist zone and regional neogene–pleistocene uplift resulted in erosion of the mesozoic sediments (gry 1969; michelsen & nielsen 1991, 1993; japsen 1993; petersen et al. 2003, this volume). northwards transgression from the tethys during early jurassic times due to an overall rise in sea level combined with continued subsidence of the north sea area and the danish basin, created an extensive epicontinental sea in northwest europe (hallam 1988, 1992; ziegler 1988). the north-eastern margin of the sea coincided with the fennoscandian border zone. coastal and delta plains were established along the margin of the sea, and hettangian–sinemurian coal-bearing strata were deposited on bornholm (rønne formation), in the øresund area, skåne (helsingborg member), and in the kattegat and north jylland (gassum formation; troedsson 1951; gry 1969; bertelsen 1978; gravesen et al. 1982; pieńkowski 1991; petersen 1993, 1994; nielsen 1995; surlyk et al. 1995). fully marine conditions were established across the entire area in the early pliensbachian, but due to a late pliensbachian regression, deposition of coal-bearing strata (sorthat formation) was resumed on bornholm until an early toarcian transgression terminated peat formation (fig. 2–3; koppelhus & nielsen 1994; petersen & nielsen 1995). in the danish basin, deposition of marine clays and sands continued until the aalenian (michelsen 1989; michelsen & nielsen 1991; nielsen 1995). continental conditions were re-established during the middle jurassic and coal-bearing strata were deposited in lake-dominated peatlands and small alluvial fans on bornholm (bagå formation; gry 1969; gravesen et al. 1982; petersen 1993; koppelhus & nielsen 1994), in lake-dominated peatlands in the øresund area (petersen 1994), and in deltaic environments in skåne (fuglunda member; rolle et al. 1979; norling et al. 1993). thin coaly muds, lacustrine clays and fluvial sands were deposited in north jylland (haldager sand formation; koch 1983; nielsen 1995). u m l u l u l u m l u l u l ø re su nd -7 ø re su nd -5 ø re su nd -1 5 ø re su nd -1 8 ø re su nd -1 3 m ar ie da l f m a f vilhelmsfält fm fortuna marl glass sand mb fuglunda mb r ya f m h ög an äs fm helsingborg mb döshult mb pankarp mb katslösa mb rydebäck mb röddinge fm ? sorthat fm bagå fm hasle fm galgeløkke mb sose bugt mb munkerup mb r øn ne f m gassum fm fjerritslev fm haldager sand fm hettangian sinemurian pliensbachian toarcian aalenian bajocian bathonian callovian (part) m id dl e lo w er ju ra ss ic system stage sw ne nw se danish basin bornholm onshore øresund wells (the measured parts) skåne ? ?? ? fig. 2. the lower and middle jurassic stratigraphy in the study area; af, annero fm. compiled from michelsen (1978), gravesen et al. (1982), sivhed (1984), ahlberg et al. (2003, this volume) and michelsen et al. (2003, this volume). note that the age range of the øresund-5 and -7 wells is poorly constrained. 634 early–middle jurassic climate establishment of peat-forming mires and preservation of peat are mainly controlled by a complex interaction of climate, subsidence and eustasy. the breakup of pangea and northwards drift of laurasia during the mesozoic into a warm-temperate to subtropical climatic zone at approximately 40°n, and the creation of a large epicontinental sea in northern europe in the jurassic, resulted in a substantial increase in humidity and rainfall compared to the hot and arid conditions that had dominated the permian and most of the triassic periods (parrish et al. 1982; ziegler 1982; hallam 1984; scotese 1994). the interaction of precipitation and temperature, and in particular seasonality of rainfall, has a strong influence on peat formation, because probably the single most important factor controlling peat distribution is continuity in rainfall (ziegler et al. 1987; calder & gibling 1994; lottes & ziegler 1994). seasonally wet conditions north of tethys (western europe) in jurassic times has been suggested by hallam (1985), and reconstructed rainfall maps for the pliensbachian suggest a moderate rainfall in the danish region (parrish et al. 1982). plant fossil assemblages from the rhaetian and lower jurassic of sweden, germany, bornholm and jameson land, east greenland, indicate that the climate was sufficiently uniform to establish a northern floral province with relatively uniform plant communities (harris 1937; florin 1958). warm conditions are suggested by the large number of plant species from jameson land, and thin cutinised leaves of podozamites and equisetales comparable in size to modern subtropical bamboos are thought to reflect favourable conditions for plant growth (harris 1937). investigations of clay mineralogy and palaeosols of the jurassic on bornholm indicate a warm, humid climate (graff-petersen & bondam 1963; arndorff 1993), and the taxonomic composition of miospore assemblages from the hettangian–sinemurian sose bugt section on bornholm lends support to the interpretation rønne– hasle fault block nyker block rønne graben arnager–sose fault block gudhjem trough bornholm high 5 6 3 4 2 1 10 km cretaceous lower–middle jurassic sorthat formation and middle jurassic bagå formation lower jurassic hasle formation lower jurassic rønne formation upper triassic kågeröd formation lower palaeozoic precambrian crystalline basement fault n fig. 3. geological map of bornholm showing the location of the studied localities: 1, coastal exposures, munkerup member; 2, sose bugt succession, sose bugt member; 3, galgeløkke cliff succession, galgeløkke member; 4, korsodde coastal cliff, sorthat formation; 5, levka-1 well, sorthat formation; 6, hasle klinkerfabrik clay pit, bagå formation. modified from gravesen et al. (1982). 635 lake & swamp lake lagoon shallow submerged inter-fluvial areas crevasse delta shoreface swamp & lake lake lagoon lagoon fluvial channel 20 m 30 40 50 60 70 80 90 seam 15 seam 12/13 seam 9 leaves branch stem coal particles fragm ents fossils iron concretions pyritic concretions carbonaceous detritus lithology mudstone siltstone silty sandstone sandstone coal biogenic structures degree of bioturbation teichichnus diplocraterion equilibrichnia horizontal burrows rootlets sedimentary structures erosional surface horizontal bedding/ lamination cross-bedding with mud drapes cross-lamination wave ripples low-angle cross-bedding small-scale hummocky cross-stratification flaser and wavy bedding lenticular and silt-streaked bedding load structure small-scale water escape structures synaeresis cracks disturbed bedding – lagoon y y y shallow submerged inter-fluvial areas three stacked fluvial channel fills three stacked fluvial channel fills stacked crevasse deltas & swamps fluvial channel crevasse delta lake & swamp lake fluvial channel lagoon m 120 130 140 150 160 clay gamma-ray gamma-ray si sand 2 4 8 mm clay si sand 2 4 8 mm levka-1 y y yy y y y yy y yyyy yy y yyyy y y yy y swamp lagoon seam 6 seam 4/5 seam 2 seam 3 110 100 fig. 4. sedimentological core-log of the upper pliensbachian – lower toarcian part of the sorthat formation in the levka-1 well (reference level: metres below surface). the investigated coal seams are indicated. the accompanying legend refers also to figs 5, 10, 11, 14, 15. intervals with no core recovery (lithology shown only) are interpreted on the basis of gamma-log data. modified from koppelhus & nielsen (1994). 636 of a wetter climate, in particular compared to areas further to the west, for example around the british isles (batten et al. 1994). however, seasonal variations are indicated by annual rings in tree trunks from the rhaetian of skåne (ahlberg & arndorff 1994), in gymnospermous wood from jameson land (harris 1937), and in tree-stumps from the hettangian and bajocian–bathonian on bornholm (höhne 1933; nielsen 1995; surlyk et al. 1995). parts of the hettangian–sinemurian succession in the øresund-18 and -15 wells are characterised by thinly-stratified light grey silty claystones and dark grey claystones or thin graded layers with a silty base and a darker, more clayey upper part. the thin layering may reflect seasonal fluctuations in sedimentation rates, possibly due to unevenly distributed rainfall through the year, as suggested by hamilton & tadros (1994) for similar stratified lacustrine mudstones with coal seams in the gunnedah basin, australia. lower jurassic coal-bearing strata sorthat formation, bornholm fully marine conditions prevailed in early jurassic times over most of the danish basin (michelsen 1975). on bornholm, however, coastal plain environments lasted until the late sinemurian; a transgression caused deposition of marine sandstones in the early pliensbachian (gravesen et al. 1982; surlyk et al. 1995). a late pliensbachian regression re-established a coastal plain environment with peat-forming mires on bornholm until the regional early toarcian sea-level rise caused a renewed transgression (koppelhus & nielsen 1994). the upper pliensbachian – lower toarcian sediments of the levka-1 well and the lower part of the korsodde section (included in the sorthat formation) were deposited in an overall coastal plain environment (figs 3–5; koppelhus & nielsen 1994; petersen & nielsen 1995; michelsen et al. 2003, this volume). the levka-1 well section represents fluvial channels, floodplain areas with shallow lakes and lagoons, and small crevasse deltas (fig. 4). coalified wood fragments and stems occur frequently. coal seams 2, 3, 6 and 9 represent peat-forming mires established on top of abandoned sandy channel fills and on heavily rooted crevasse and lake deposits in shallow inter-fluvial areas. seams 4 and 5 represent peat accumulation on top of lagoonal sediments showing wave-ripples, wavy and flaser bedding, bioturbation and transported equisetites stems. a well-oxidised palaeosol that developed upon the fill of a coastal lake, is overlain by seams 12 and 13 which are separated by a thin sandy, rooted parting. the clay of the lake-fill contains a palynomorph assemblage dominated by spores and pollen, together with a few acritarchs and tasmanites. seam 15 is interpreted to represent an infilled-lake peat mire. the lower 133 m of the korsodde section consists of fluvial channel sand interbedded with heterolithic fining-upwards units of clay and silt and laminated to seam 4 seam 3 clay si sand 8 mm42 86 90 92 m 88 lakes & swamps 118 120 122 124 m seam 5 estuarine channel lagoon lagoon korsodde section clay si sand 8 mm42 fig. 5. sedimentological log of selected coal-bearing intervals from the korsodde coastal cliff succession (sorthat formation); reference level: metres above base of section. modified from koppelhus & nielsen (1994); for legend, see fig. 4. 637 almost massive clay beds topped by coal seams that overlie rooted horizons (fig. 5). the fining-upwards units contain pyrite nodules, coalified wood fragments, and mixed palynomorph assemblages of spores, pollen, acritarchs, freshwater algae and marine dinoflagellate cysts, and were deposited in small coastal lakes and protected lagoons (batten et al. 1994; koppelhus & nielsen 1994). coal seam composition and peat-forming environments the studied coal seams of the sorthat formation are between 0.08 and 0.57 m thick (figs 4, 5). those from the levka-1 well and seam 4 from the korsodde section are dominated by very high contents of huminite (seam-average ranges from 80–90 vol.%), and only minor amounts of inertinite, liptinite and minerals (petersen & nielsen 1995). pyrite occurs in several of the seams. seams 3 and 5 from the korsodde section differ in petrographic composition from the rest of the investigated seams. seam 3 contains significant amounts of minerals (seam-average 37 vol.%) and seam 5 contains a high proportion of inertinite (seam-average 30 vol.%). both seams contain pyrite. reflected-light microscopic analysis of a sample from seam 2 in levka-1 reveals that it contains 2 vol.% pyrite and 6 vol.% of other minerals (petersen & nielsen 1995). the composition of the inorganic components in the sample was further investigated by computer controlled scanning electron microscopy (ccsem). inorganic parunclassified rutile gypsum/ al-silicate gypsum quartz kaolinite illite quartz/ pyrite clay/pyrite pyrite 0 10 20 30 40 50 wt% (mineral basis) m in er al s sorthat fm, seam 2 bagå fm, seam 6 fig. 6. composition (wt% on a mineral basis) of the mineral matter in the upper part of seam 2, sorthat formation (levka-1 well) and in seam 6 (interval 125–135 cm) in the upper bagå formation. the group ‘unclassified’ includes minerals which constitute less than 1 wt% of the inorganic fraction. lower delta plain upper delta plain strand plain piedmont plain back barrier tree density increases 0.0 1.0 2.0 3.0 4.0 0.1 10 1 tissue preservation index g el ifi ca tio n in de x lower jurassic : sorthat formation; levka-1 well and korsodde coastal cliff : rønne formation, galgeløkke member; galgeløkke coastal cliff : øresund area; øresund-13, -15 and -18 wells 100 fig. 7. coal samples from the lower jurassic of bornholm and the øresund area plotted on the tissue preservation index (tpi) vs. gelification index (gi) diagram of diessel (1986). the strand plain area is after kalkreuth & leckie (1989; see also diessel 1992). notable are the low gi values of the galgeløkke member coal suggesting oxidizing conditions during deposition of the peat (domed bog), and the very high gi values of the upper pliensbachian – lower toarcian sorthat formation coals suggesting anoxic, waterlogged conditions during peat deposition. 638 fig. 8. colour photomicrographs of pyrite (reflected white light, oil immersion, scale bar c. 30 µm, except a2: scale bar c. 20 µm). a1: framboidal pyrite (p) in ulminite in coal seam 2, sorthat formation. note how the sporinite (arrows) ‘flows’ around the pyrite, suggesting precipitation of the pyrite (syngenetic) prior to compaction (seam 2, sample 291934; levka-1 well). a2: framboidal pyrite (p) in the same sample. b: pyrite (p) filling cleats in ulminite in seam 8 in the upper bagå formation, suggesting epigenetic pyrite precipitated after gelification of the organic matter (seam 8, sample b-9.2, hasle klinkerfabrik clay pit). c: pyrite (p) filling cleats in macrinite in seam 8 in the upper bagå formation, suggesting epigenetic pyrite precipitated after gelification and oxidation of the organic matter (seam 8, sample b-9.2, hasle klinkerfabrik clay pit). d: pyrite (p) precipitated in cell lumens in fusinite in seam 5, upper part of the bagå formation (seam 5, sample b-6.3, hasle klinkerfabrik clay pit). a1 p p p p pp b d c a2 ticles below 1 µm and organically bound elements are not detected. the ccsem analysis indicates that the inorganic fraction is dominated by 45 wt% quartz with subordinate pyrite and minor proportions of gypsum and clay minerals (fig. 6). in general, the quartz is in the size range of clay and silt with the major part being fine-grained silt (4–16 µm). the pyrite crystals are mainly within the size limits of silt followed by very fine-grained and fine-grained sand. framboidal aggregates may account for the larger particles. the peats accumulated mainly in anoxic and fully water-saturated freshwater mires as indicated by the very high content of humified organic matter in all of the seams in the levka-1 section (high gelification index values, gi = huminite/inertinite, fig. 7; petersen & nielsen 1995). the high content of humocollinite and the abundance of hopanoids in the coals indicate pronounced bacterial activity reflecting abundance of nutrients and a relatively high ph level (teichmüller 1989; waples & machihara 1991; peters & moldowan 1993). the significant amounts of pyrite in seams 2 and 4 and minor amounts in several other seams suggest occasional brackish – marine influence on the mires (cohen et al. 1984; phillips & bustin 1996a). the pyrite occurs in association with fusinite and inertinite, and as framboidal aggregates and minute crystals in huminite (fig. 8a). it is interpreted to be mainly syngenetic. however, some of the pyrite, e.g. in seam 4, occurs infilling cleats in huminite and macrinite, thus suggesting that precipitation of pyrite also took place after gelification. the very small grain size of the mineral matter (clay to fine-grained silt) suggests that the mineral matter was deposited from suspension or may have been windborne. three coal seams, seams 3, 4 and 5, from the korsodde succession have been investigated in detail (fig. 5; petersen & nielsen 1995). seam 4 is petrographically very similar to the seams from the levka-1 well and represents a wet, anoxic, and probably rheotrophic, nutrient-rich peat-forming environment. the seam overlies sediments from a shallow, vegetated coastal lake, that occasionally was flooded by the sea. seam 3 also overlies coastal lake deposits, and the high mineral content in the seam reflects the development of the lake into a frequently inundated open mire. a peat mire failed to develop although rootlets indicate that in situ plants contributed to peat accumulation. seam 5 overlies a lagoonal unit and constitutes the final, regressive phase in the evolution of an estuarine complex. the precursor peat was subjected to desiccation, possibly linked to a decreased rate of watertable rise. palynological evidence of the peat-forming vegetation palynological analyses were carried out on five coal samples from seams 2 and 15 in levka-1 and three samples from the sediment immediately below the coal seams. although the parent plant affinities of palynomorphs from the jurassic are uncertain in most cases, it has been possible to discriminate several groups of plants: ferns (pteridophytes) of the order filicales (class: filicopsida), cycadophytes of the order cycadales (class: gymnospermopsida), coniferophytes of the order coniferales (class: gymnospermopsida), mosses (division: bryophyta), and microplankton-like acritarchs, marine green-algae (tasmanites spp.) and a freshwater greenalgae (botryococcus spp.; table 1). the palynomorph assemblage of the sediment sample below seam 2 in the levka-1 well is dominated by cycadalean pollen (c. 25%, chasmatosporites hians, c. minor, c. apertus) and jurassic species of the family pinaceae (c. 50%, mainly cerebropollenites thiergartii) (table 1). the content of ferns is likewise pronounced (18%). moving up to the two coal samples, several pollen types that have affinities with conifers disappear, while chasmatosporites hians, whose affinity lies with the order cycadales (van konijnenburg-van cittert 1971), and cerebropollenites thiergartii, whose affinity lies with the family pinaceae, increase significantly in number upwards in the seam (cycadales and pinaceae pollen increases to 14% and 19% in the uppermost sample, respectively). however, the pollen corollina torosus, the affinity of which is considered to lie with the large conifer cheirolepidaceae (srivastava 1976) and which constituted an important part of the jurassic coniferous forests (stewart 1983), appears in the lower coal sample. the occurrence of these larger gymnospermous plants is in agreement with the considerable content of the maceral subgroup humotelinite in the lower coal sample. the diversity of spores whose affinities lie with the ferns, in particular the family osmundaceae, increases from the underlying sediment up into the coal. in the lower coal sample, the proportion of ferns reaches 56%, suggesting a mire with a pronounced herbaceous vegetation. the generally high content of degraded organic matter (humodetrinite) in both coal samples correlates with a high proportion of these small plants. spores related to mosses constitute a significant proportion of the coal assemblages. the palynomorph diversity is lower in the uppermost coal sample, possibly due to the influence of marine water as interpreted from the coal petrology and the occur639 640 locality seam sample sample type filicopsida % lycopsida % cycadales gymnosperms % coniferales p* t† g‡ microplankton§ % bryophyta % n� øresund-13 øresund-18 levka-1 bagå clay pit 1 2 cm below seam 1 4 cm above seam 1 1 1 1 1 2 cm below seam 1 4 cm below seam 1 2 2 2 cm below seam 2 2 2 3 cm below seam 2 15 15 15 2 cm below seam 15 3 cm below seam 15 1 2 3 3 4 5 6 few cm below 6 247943 6371 6373 6372 6369 247944 247945 6375 6374 247947 clay/silt coal coal coal coal light clay light clay coal coal dark clay 5.1 16.0 23.7 41.7 0 73.9 71.4 61.4 46.0 20.3 6.2 2.5 1.7 0 0 0 0 2.6 9.8 8.8 1.1 7.3 5.1 16.7 0 4.3 4.4 1.3 2.0 14.7 10.7 14.1 8.5 0 40.0 0 1.1 18.6 6.9 10.4 14.1 25.6 30.5 8.3 0 0 3.3 1.3 2.0 4.7 4.6 0.5 1.7 12.5 ¶ 0 14.4 4.4 0 1.0 5.1 0 0.5 5.1 4.2 20.0 1.5 4.4 0 0 0 178 206 59 24 5 69 91 75 102 193 291934 291932 247938 6386 6383 6379 247940 247941 coal coal coaly clay coal coal coal clay clay 10.8 56.1 18.0 33.3 24.5 28.9 43.5 43.4 13.5 8.4 24.6 28.5 40.0 42.0 11.9 18.2 18.9 6.3 50.2 12.4 11.1 8.1 33.7 31.5 0 6.3 0.6 4.9 3.3 1.6 2.0 0.5 32.5 12.5 6.6 19.7 14.4 11.3 8.9 6.4 5.4 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 18.9 10.4 0 1.2 6.7 8.1 0 0 37 48 167 81 90 62 202 203 b-1.1 b-2.1 b-3.2 b-3.1 b-4.1 b-6.9 b-7.2 b-7.1 coal coal coal coal coal coal coal black clay 62.7 31.3 50.9 54.1 53.3 24.7 30.7 36.2 13.2 9.7 10.5 16.2 22.3 31.1 3.6 21.3 14.2 25.7 12.3 2.7 22.2 24.6 19.8 27.5 0 8.3 0 2.7 0 0 0 0 5.5 20.8 24.6 21.6 2.2 19.6 45.9 12.5 0 0 0 2.7 0 0 0 0 0 0 0 0 0 0 0 0 4.4 4.2 1.7 0 0 0 0 2.5 91 144 57 37 45 61 111 80 6368 247948 64.3 14.0 5.1 4.2 3.5 13.5 11.2 19.2 0.5 8.2 15.4 37.9 0 3.0 0 0 196 209 table 1. peat-forming plants based on palynological evidence coal black clay ¶ three dinoflagellates. �number of palynomorphs counted. ‡ other gymnospermous pollen. § mainly acritarchs, some tasmanites and botryococcus. * pinaceae. † taxodiaceae. 58.2 33.5 23.7 16.6 40.0 5.9 11.0 14.8 32.3 36.0 0 10 20 30 40 50 60 70 80 90 100 27.10 27.20 27.30 27.40 27.50 m levka-1 well, seam 15 27.60 81 90 n 62 202 203 ferns (filicopsida) cycadophytes (cycadales) coniferophytes (coniferales) coal interpreted plant groups lithology mudstone gymnosperms, other mosses (bryophyta) % (cumulative) fig. 9. the succession of main plant groups in seam 15, levka-1 well, as suggested by the botanical affinities of palynomorphs and their abundance in percentages. n, number of palynomorphs counted. 641 rence of a probable acritarch species (lecaniella foveata; batten et al. 1994; petersen & nielsen 1995). the clay samples taken immediately below seam 15 in the levka-1 well contain a significant content of the same pollen species identified in seam 2 whose affinities lie with the order cycadales and jurassic species of the family pinaceae (fig. 9; table 1). fern spores are present below seam 15 in considerable amounts (43%, mainly deltoidospora spp.), and the diversity of spores related to ferns is greater than in the sediment sample below seam 2. the most pronounced changes in the palynomorph assemblages in seam 15 compared to the underlying sediment are a significant decrease in the proportion of pollen whose affinities are with the family pinaceae, a decrease in the amount of fern spores, and a considerable increase in the proportion of pollen whose affinities are with the cycadophytes (fig. 9; table 1). despite differences in types and amounts of the recorded pollen species among the three coal samples in seam 15, they display a rather uniform composition. the coal petrographic composition is also uniform, conforming with studies on recent peat deposits, which have shown a correlation between peat vegetation and peat composition (cohen & spackman 1977; phillips & bustin 1996b). thus, the mire vegetation may have been dominated by gymnospermous plants and a secondary proportion of ferns characterised by the genera dicksonia or coniopteris and the family osmundaceae (tralau 1968; van konijnenburg-van cittert 1978). this might correlate with the dominance of humotelinite over humodetrinite in the seam. the lowermost coal sample has both the smallest amount of gymnospermous pollen and the lowest proportion of humotelinite of the three coal samples. in general, the palynomorph assemblages suggest that the peat-forming vegetation of seam 15 was of higher diversity than that of seam 2, which may be explained by the influence of saline water in the precursor mire of seam 2. duration of peat formation the compaction ratio for peat to bituminous coal varies between 4.1:1 and 30:1 (ryer & langer 1980). a peat:coal compaction ratio of 10:1 may be assumed for the subbituminous coals in this study (see also mccabe 1991). the tropical batang hari river peat deposit in sumatra has an accumulation rate of c. 1.8 mm/yr, whereas the cold temperate peat deposit in maine has been accumulating at a rate of c. 0.6 mm/yr (cameron et al. 1989). an average accumulation rate of 1 mm/yr may thus be a reasonable assumption for the warm temperate to subtropical jurassic climate. using this assumption, the late pliensbachian – early toarcian peat-forming mires on bornholm existed between 800 and 5700 years, with an average of c. 2300 years. other lower jurassic coal-bearing strata in the fennoscandian border zone hettangian–sinemurian mires are represented by coal and coaly mudstones within the munkerup member and the sose bugt member of the arnager–sose fault block on the south coast of bornholm (figs 2, 3). the poorly exposed munkerup member consists primarily of lake deposits (gry 1969; gravesen et al. 1982; koppelhus 1991). periodically, the lacustrine environment developed into mires. increasing deoxygenation of the lakes and encroachment of vegetation is evident sose bugt member galgeløkke member rønne formation m 10 5 0 seam a seam b seam c seam d m 25 20 15 claysi sand claysi sand fig. 10. sedimentological log of the coastal sose bugt succession (hettangian–sinemurian sose bugt member) and galgeløkke cliff section (sinemurian galgeløkke member). after gravesen et al. (1982); section heights refer to the original sections of gravesen et al. (1982). for legend, see fig. 4. from the gradual transition from clayey, rooted sediments to the overlying coaly mudstones. the coaly mudstones are inertinite-rich, contain pyrite, and are to a large extent allochthonous, reflecting deposition in freshwater to brackish, open water environments (petersen 1993). the sose bugt member consists of lacustrine, fluvial, lagoonal, and restricted marine shoreface to offshore sediments (figs 2, 3, 10; surlyk et al. 1995). the occurrence of marine palynomorphs at certain levels within the terrestrial part of the succession indicates occasional marine incursions. the four 0.12–0.30 m thick, mainly allochthonous, mineral-rich coaly mudstones were deposited in shallow lakes on a coastal plain (petersen 1993; surlyk et al. 1995). the mires developed largely under freshwater conditions, but may have periodically been influenced by marine water. in the neighbouring rønne–hasle fault block, the rønne formation is dominated by lagoonal units with coal seams deposited in an overall coastal plain environment (nielsen 1995). the exposed section of the sinemurian galgeløkke member consists of sandy and muddy lagoonal, tidal flat and tidal channel complex deposits (figs 3, 10; sellwood 1972; gravesen et al. 1982). an inertinite-rich coal seam, 0.10–0.35 m thick, overlies wavyand flaser-bedded tidal flat heteroliths. the seam represents a dominantly freshwater peat bog 642 lagoon or restricted brackish bay lagoon or restricted brackish bay sparsely vegetated open water environment. limnic facies. mainly allochthonous deposition. a øresund-13 well m 88 clay silt sand 89 90 91 92 93 94 upper part: densely vegetated rheotrophic infilled lake swamp. telmatic freshwater facies. lower part: sparsely vegetated open water lake. limnic facies. mainly allochthonous deposition. freshwater lake upper part: mesotrophic to ombrotrophic domed bog. (telmatic to) terrestrial freshwater facies. lowermost part: densely vegetated rheotrophic infilled lake swamp. telmatic freshwater facies. freshwater lake b øresund-18 well m 10 seam 2 11 12 ? 13 14 15 16 clay seam 1 siltsilt sand densely vegetated rheotrophic swamp. slightly brackish wet telmatic facies. restricted brackish shallow water lake c øresund-15 well m 101 ? clay siltsilt sand 102 103 104 105 fig. 11. a: sedimentological core-log of the studied hettangian–sinemurian part of the øresund-13 well. b: sedimentological core-log of the studied hettangian–sinemurian part of the øresund-18 well. c: sedimentological core-log of the studied sinemurian part of the øresund-15 well. modified after petersen (1994); reference level: metres below sea level. for legend, see fig. 4. with a domed surface (fig. 7, gi values < 1), implying a significant change from the tidal flat environment that prevailed prior to peat formation (petersen 1993), and the possible presence of a diastem below the coal seam. siliciclastic deposition may have occurred contemporaneously nearby, but the elevated peat surface hindered flooding of the peat deposit (e.g. staub & esterle 1994; phillips & bustin 1996b). the high content of well-preserved pyrofusinite, which represents fossil charcoal (scott 1989), indicates that wildfires occurred in the bog. this resulted in a high inertinite content in the seam as shown by low gi values in the tpi versus gi diagram (fig. 7). the coal seam contains small amounts of pyrite towards the top, which probably was precipitated due to migration of sulphate ions into the peat from the overlying brackish sediments after peat formation was terminated. similar paralic depositional environments existed in the øresund area in hettangian–sinemurian times (fig. 1). the bioturbated heteroliths underlying a 0.18 m thick, inertiniteand mineral-rich coal seam in the øresund-13 well represent deposition in a shallow, restricted bay or lagoon (fig. 11a ; petersen 1994). after the shallow area was filled, it became vegetated and accumulation of mainly allochthonous organic matter resistant to degradation during transportation took place in an open freshwater mire. both seams 1 (0.25 m) and 2 (0.15 m) in the øresund-18 well overlie lake fill sediments (fig. 11b; petersen 1994). the presence of rooted claystones below the seams demonstrates that the gradual fill of the lakes promoted plant growth. both seams represent rheotrophic, probably nutrient-rich freshwater mires, although the precursor peat of seam 1 may temporarily have evolved to a more domed state. sinemurian coal-bearing strata occur in the øresund-15 well, and the claystones underlying the coal seam were deposited in a shallow, weakly brackish lake situated in a coastal plain (fig. 11c). the preserved stratification and absence of trace fossils suggest a poorly oxygenated sediment. the 0.16 m thick, huminite-dominated coal seam represents an occasionally slightly brackish, nutrient-rich swamp established on top of the lake-fill. a c. 3 cm thick coaly mudstone occurs in the uppermost lower jurassic in the skagen-2 well in north jylland (referred to as the fjerritslev formation by michelsen 1978). the coaly mudstone represents an allochthonous deposit, which is emphasised by the high content of the mineral-rich microlithotype carbominerite (71 vol.%), microlaminated detrital macerals, and macerals resistant to degradation during transportation (inertinite: 39 vol.%; liptinite: 9 vol.%). oxidised cutinites may reflect oxidation of the organic matter during transportation. the occurrence of pyrite (2 vol.%) suggests that the open water environment was brackish or periodically influenced by marine water. the peat mire was situated in a coastal environment, consistent with the position of the bed within a thin unit that correlates with transgressive marine mudstones further to the west (i.e. basinwards; nielsen 1995). palynological evidence of the peat-forming vegetation in the øresund area a total of twelve samples were analysed from the øresund-13 and -18 wells; these comprise seven coal samples, four samples from the sediment immediately below the coal seams, and one sample immediately above a seam. in addition to the plant groups mentioned from the sorthat formation coals, club mosses (class: lycopsida) were identified in these coals (table 1). a sample was taken from the coal seam in the øresund-13 well and from the black clay about 2 cm below the seam. the latter shows a higher diversity than the coal sample. acritarchs (lecaniella foveata, leiosphaeridia spp. and a possible cymatiosphaera spp.), tasmanites spp. and a possible dinoflagellate cyst (? mancodinium semitabulatum) occur only in the sediment sample. this is consistent with the interpretation of the sedimentary environment as a lagoon or brackish restricted bay and the seam as a freshwater deposit (petersen 1994). spores related to ferns, in particular deltoidospora spp., whose affinities lie with the two fern genera coniopteris or dicksonia (tralau 1968), are common in the clay. however, coniferalean and cycadalean pollen are more abundant. bisaccate species and a pollen of unknown affinity are also abundant. in combination, the pollen related to gymnosperms constitute 78% of the palynomorph assemblage in the clay sample (table 1). in contrast, in the coal sample the proportion of gymnosperm pollen is only 31%. in particular, the proportions of the species chasmatosporites hians and perinopollenites elatoides, whose affinities are with the cycadales and the coniferophyte family taxodiaceae (guy-ohlsson 1986) respectively, are significantly lower. however, spores whose affinities lie with the order lycopodiales, the fern genera dicksonia and the fern family osmundaceae (van konijnenburg-van cittert 1978) are more abundant. in total, the proportion of fernrelated spores increases to approximately 64% in the coal sample. thus, the palynomorphs indicate a change 643 644 in the composition of the vegetation from the siliciclastic environment to the mire environment. the sparse mire vegetation probably consisted to a large extent of herbaceous plants. the sediment samples below seam 1 in the øresund-18 well are dominated by fern spores (71–74%, mainly deltoidospora spp.) and up to 15% acritarchs (fig. 12; table 1). upwards through the coal seam, spores related to ferns decrease and constitute only 16% in the uppermost sample. club mosses (retitriletes clavatoides) were probably present in the late peat stages corresponding to slightly wetter conditions. most notable compared to the underlying clay is the significant content of taxodiaceous pollen (perinopollenites elatoides) and bisaccate pollen related to unknown gymnosperm plants. pollen and spores also show that plants belonging to the cycadales (chasmatosporites hians, c. apertus), the pinaceae (pinuspollenites minimus, cerebropollenites macroverrucosus, c. thiergartii), and the bryophyta constituted parts of the original mire vegetation (fig. 12; table 1). an evolutionary link between the parent plant of c. macroverrucosus and the extant tsuga has recently been suggested, although the parent plant of the former probably became extinct during the cretaceous (batten & dutta 1997). acritarchs and dinoflagellates are present. the composition of the vegetation cannot be directly correlated to the maceral composition of the seam, which may be due to the high proportion of oxidised organic matter (inertinite), particularly inertodetrinite. the occurrence of club mosses in the upper part of the seam may correspond to a return to more wet and nutrient-rich conditions, which would be consistent with the increased content of mineral matter. in contrast, the occurrence of microplankton in the seam is of particular interest. no indications of saline water influence, such as pyrite, have been observed in the seam, and the sediments have been interpreted as freshwater deposits (petersen 1994). however, the occurrence of acritarchs suggests, at least temporarily, brackish conditions. on the basis of the palynomorph assemblage, the mire may have been covered with a dominantly gymnospermous vegetation and a significant, subordinate proportion of ferns and club mosses. spores (mainly deltoidospora spp.) probably derived from ferns increase from c. 20% in the dark clay immediately below seam 2 in the øresund-18 well to 61% in the upper part of the seam, whereas the proportions of cycadalean pollen and pollen with unknown gymnosperm affinities decrease significantly upwards (table 1). the contribution of the taxodiaceae family was much less important in the precursor mire of seam 2 than in that of seam 1. the mire may have been characterised by a prominent herbaceous vegetation, and the highest proportion of spores derived from club mosses occurs in the lower part of the seam representing wet, 12.00 12.10 12.20 12.30 12.40 m 178 206 59 24 69 91 n øresund-18 well, seam 1 0 10 20 30 40 50 60 70 80 90 100 % (cumulative) ferns (filicopsida) club mosses (lycopsida) cycadophytes (cycadales) coniferophytes (coniferales) coal interpreted plant groups lithology mudstone siltstone carbonaceous detritus gymnosperms, other microplankton (mainly acritarchs) mosses (bryophyta) fig. 12. the succession of main plant groups in seam 1, øresund-18 well, as suggested by the botanical affinities of palynomorphs and their abundance in percentages. n, number of palynomorphs counted. open-water conditions. the contribution of plant tissue from ferns together with club mosses in the precursor mire is compatible with the high content of detrital organic matter (humodetrinite) in the coal. duration of peat formation the coaly mudstones from the sose bugt member are very low in rank (petersen et al. 2003, this volume), and a compaction ratio of 4:1 for peat:lignite (esterle & ferm 1994) may provide a better estimate than that adopted for the sub-bituminous coals (see above). using this value and ignoring the relatively high content of mineral matter, the original carbonaceous deposits were between 0.5 and 1.2 m thick, and organic matter accumulation lasted for about 500–1200 years. the hettangian munkerup member is located in the same fault block as the hettangian–sinemurian sose bugt member. the boundary between the two members is not exposed (gry 1969; gravesen et al. 1982), but the thickness of the sediment package between the seams of the two members is insignificant with regard to coalification. thus, although the munkerup member coals yield higher reflectance values than the sose bugt member seams (petersen et al. 2003, this volume), the 4:1 compaction ratio is also applied to the munkerup member coals. using this compaction ratio, the original peat deposits of the munkerup member were also between 0.5 and 1.2 m thick, and peat accumulation lasted for about 500–1200 years. the original peat deposit represented by the coal seam in the galgeløkke coastal cliff section in the rønne graben had a maximum thickness of 3.5 m (peat:coal compaction ratio: 10:1), corresponding to 3500 years of peat accumulation. on average, the mires represented by the investigated seams in the rønne formation thus existed for c. 1200 years. assuming a 10:1 compaction ratio for the lower jurassic coal seams in the øresund area, the original peat thickness was 1.5–2.5 m and the duration of peat accumulation was 1500–2500 years. on average, the peat mires existed for 1850 years. middle jurassic coal-bearing strata bagå formation, bornholm the lower to middle part of the bagå formation in the hasle klinkerfabrik clay pit is a heterogeneous, clayrich succession about 60 m thick that is characterised by stacked units (c. 10 m thick) of laminated, grey – dark grey clay with silt and sand laminae and subordinate beds of cross-bedded or laminated sandstones, often showing rootlets (koppelhus & nielsen 1994). these units are separated by coaly mudstones or coal seams, associated with rooted horizons that record the periodic establishment of peat-forming mires. fossil plant fragments, small logs, stems, and leaves are common in the clay. the palynomorph assemblages are all non-marine and dominated by pteridophyte spores and gymnosperm pollen (hoelstad 1985; koppelhus & nielsen 1994), and the total sulphur content in the coaly mudstones and coal seams is generally less than 0.80 wt%, indicating a freshwater environment (casagrande 1987). thus, the clay was deposited in freshwater lakes, which were gradually filled and transformed into openwater or peat-forming freshwater mires. the sand beds represent crevasse splays and channel deposits. the upper part of the bagå formation becomes more sandy, with some very poorly sorted sand beds, occasionally with large kaolinised granite boulders deposited by debris flows that originated from the granite horst located immediately to the east (gry 1969; gravesen et al. 1982; nielsen 1995). pyrite nodules are common in the sand beds. based on the clastic facies and the proximity to the major graben fault against the granitic basement, a small alluvial fan environment with deposition from sheet wash and mud flows is envisaged. coal seam composition and peat-forming environments eight coal seams, 0.05–1.90 m thick, have been investigated in the upper part of the bagå formation (seams 1–8 of petersen 1993). the seams are petrographically complex, with varying amounts of inertinite (1–72 vol.%; 74% of the studied coal samples contain 20–60 vol.%) and huminite (16–89 vol.%; 71% of the studied samples contain more than 30 vol.%), and often a considerable content of minerals (2–57 vol.%; 57% of the studied samples contain 10–20 vol.%; petersen 1993). fusinite is a prominent component of the inertinite maceral group (up to 41 vol.%). pyrite occurs in seven of the eight seams. framboidal pyrite has not been observed, although in several of the seams significant amounts of pyrite (up to 36 vol.%) fill cell-lumens of fusinite, occur as massive pyrite or fill cleats in gelified organic matter (fig. 8b, c, d; petersen 1993). analysis of a sample from seam 5 by ccsem shows that the inorganic fraction is dominated by pyrite (46 wt%), 645 646 followed by quartz (19 wt%), clay associated with pyrite (13 wt%) and kaolinite (10 wt%; fig. 6). the majority of the quartz is in the grain-size range of silt (in particular coarse silt, 32–64 µm), to fine-grained sand, whereas most of the pyrite is of fine-grained sand grade. the eight coal seams from the upper part of the bagå formation were formed in freshwater swamps, that occasionally were established on the alluvial plain. the occurrence of several well-defined root horizons below a thick coal seam (seam 5 of petersen 1993) indicates repeated episodes of siliciclastic deposition followed by colonisation of the sediments by plants. a level coloured by humus substances below each root horizon was possibly formed by a low-standing watertable, which together with the petrography of seam 5 imply an unstable watertable. most of the samples from seam 5 plot in the field of the piedmont plain in the tpi versus gi diagram (fig. 13), consistent with the alluvial fan interpretation. varying amounts of detrital minerals in the seams is attributed to periodic outwash from the nearby granitic horst. in addition to pyrite and clay associated with pyrite, the inorganic fraction in the 125–135 cm interval in seam 5 is characterised by quartz, primarily of coarse silt to fine-grained sand size, and kaolinite. the comparatively high content of kaolinite is probably related to the proximity of the mires to the weathered granitic basement of the bornholm high to the east. a large proportion of impure coals and an abundance of inertinite-rich lithotypes occur in piedmont coal seams in the springhill coalfield of nova scotia, where peatlands were nourished by groundwater discharge from alluvial fans and basin-margin fault-fed springs (calder 1993, 1994). a similar small-scale scenario is envisaged for the sand-dominated, upper part of the bagå formation. formation of dull coals due to severe oxidation in the piedmont plain setting has also been noted by diessel (1986), and most of the upper bagå formation coals contain much inertinite. the inertinite may have formed due to a fluctuating watertable in the precursor mires, an inertinite-forming process known from recent peat-forming environments (cohen & stack 1996; moore et al. 1996). fluctuations were probably related to seasonality as indicated by the occurrence of annual rings in wood from the coal seams in the bagå formation (höhne 1933). a peat-forming flora dominated by cellulose-rich plants (shrubs, herbs) is suggested by the maceral composition. palynological evidence for the peat-forming vegetation seven coal samples and one clay sample were analysed from coal seams 1–6 (petersen 1993) in the upper part of the bagå formation. it is possible to discriminate several plant groups: ferns (order: filicales; class: lower delta plain upper delta plain strand plain piedmont plain back barrier tree density increases 0 1.0 2.0 3.0 4.0 5.0 0.1 100 10 1 tissue preservation index g el ifi ca tio n in de x middle jurassic : bagå formation (uppermost part); hasle klinkerfabrik clay pit : øresund area; øresund-5 and -7 wells fig. 13. coal samples from the middle jurassic on bornholm and the øresund area plotted on the tpi vs. gi diagram (fig. 7). most of the samples from the upper bagå formation plot within the field of the piedmont plain. filicopsida), club mosses (class: lycopsida), coniferophytes (order: coniferales; class: gymnospermopsida) and mosses (division: bryophyta; table 1). pollen related to cycadophytes were not found, although höhne (1933) identified cycadophytes in the coals. a sample from the lowermost part of seam 1 is characterised by a high diversity of spores, which are related to ferns of the family osmundaceae and the genera dicksonia or coniopteris (tralau 1968; van konijnenburgvan cittert 1978), to the orders lycopodiales and selaginellales, and to bryophytes. fern-related spores constitute 63% and club mosses 13% of the assemblage (table 1). pollen whose affinities lie with the cheirolepidaceae and pinaceae are also present. compared to seam 1, seam 2 is characterised by a higher diversity and a larger proportion of pollen whose affinities are with the families pinaceae, taxodiaceae, and other gymnosperms (table 1). the spore assemblage has a different composition from that of seam 1 and makes up a smaller proportion of the total palynomorphs. it implies, however, a rather similar overall composition of the herbaceous vegetation. based on coal petrography, petersen (1993) suggested that the vegetation in the mires represented by seams 1 and 2 was generally small-sized and consisted of abundant herbaceous plants, small trees, shrubs and tree ferns. palynological evidence of tree ferns has not been found in the coals, and the abundance of spores whose affinities are with herbaceous plants indicates that ferns and club mosses were more common in these mires than deduced from the maceral composition, which has a rather high content of humotelinite. the rest of the samples from seams 3–6 are rather similar. the diversity of spores related to ferns is considerably lower than in seams 1 and 2; however, they still constitute up to 54% of the assemblages (table 1). ferns in particular seem to have been abundant in the mires represented by seams 3 and 4. the most abundant is deltoidospora spp., whose affinity is with the genera dicksonia or coniopteris (tralau 1968). the proportion of club mosses was significant in the mires, and the spores (mainly retitriletes austroclavatoides, r. clavatoides, r. semimuris, sestrosporites pseudoalveolatus) amount to 22% and 31% in seams 4 and 5, respectively. these spores are particularly associated with mineral-rich levels in the coals implying that the plants favoured wetter and more nutrient-rich conditions. the pollen species cerebropollenites macroverrucosus which is related to the gymnosperm genus tsuga of the family pinaceae (tralau 1968; batten & dutta, 1997), is abundant. in the coal sample from seam 6, the pollen vitreisporites pallidus is abundant. the affinity of this pollen is with the caytoniales (van konijnenburg-van cittert 1971), a gymnosperm order of uncertain relationship (stewart 1983). a peat-forming vegetation consisting of abundant herbaceous and shrubby plants and some larger trees was suggested by petersen (1993). to some extent this is supported by the palynomorph assemblages, but the proportion of gymnospermous trees may have been greater than deduced from the maceral composition. plants belonging to the family pinaceae were present. reconstruction of the gymnospermous vegetation in the mires is difficult due to the high proportion of pollen of unknown affinity (table 1). duration of peat formation assuming a peat accumulation rate of 1 mm/yr and a peat:coal compaction ratio of 10:1 (see above), the original peat deposits represented by the coal seams in the upper part of the bagå formation were between 0.5 and 19.0 m thick, and peat accumulation occurred in the mires for between 500 and 19 000 years. other middle jurassic coal-bearing strata in the fennoscandian border zone several seams, 0.16–0.70 m thick, are present in the øresund-5 and -7 wells, overlying parallel-laminated to homogeneous claystones deposited in freshwater lakes (figs 14, 15; petersen 1994). the abundance of coalified plant remains and the upwards-increasing content of organic matter, coal particles, coal streaks and rootlets in the claystones towards the coal seams record the gradual fill of the lakes and deterioration of the oxygen supply in the mires. the lower part of the øresund-5 seam represents a drier mire environment, whereas the upper part of the seam is a carbonaceous claystone deposited in an open freshwater environment. this indicates that the rising groundwater level gradually outpaced organic matter accumulation, and peat formation was terminated by lacustrine flooding. seams 1, 3 and 4 in the øresund7 well contain a higher inertinite content as shown by low gi values (fig. 13), and they represent bogs that were probably subjected to a fluctuating watertable (petersen 1994). the mire represented by seam 2 was also established after a lake had been filled with sediment. however, the mire was continuously water-saturated and anoxic, and probably nutrient-rich and rheotrophic. 647 648 the middle jurassic peat-forming freshwater mires in the øresund area were thus located in an overall lakedominated wetland with a prolific flora indicated not only by the coal seams but also by the abundance of coalified leaves and branches in the claystones. there is no evidence of marine influence, and the continental sedimentation kept pace with the subsidence-induced base-level rise. in north jylland, thin coal seams and carbonaceous mudstones are present in the middle jurassic haldager sand formation. analysis of a 4 cm thick seam in the frederikshavn-3 well shows that it is inertiniteand mineral-rich (75 vol.% and 17 vol.%, respectively), which may indicate an allochthonous origin. however, black carbonaceous clay and rootlets below the seam indicate a gradual change to oxygen deficient, low energy conditions and the establishment of vegetation. degraded liptinite components suggest fungal and/or bacterial attack under well-oxidised conditions. this and the scarcity of huminite (1 vol.%) may imply that the seam was severely oxidised, possibly due to a fall in the watertable following the accumulation of the organic matter. minerals are easily incorporated into a thin layer of organic matter, and the carbonaceous seam is probably autochthonous and was deposited in a shallow, slightly brackish area, as suggested by the presence of pyrite (4 vol.%). the seam is overlain by coaly matter associated with quartz grains, and poorly sorted darkbrown clay and sand, indicating a gradually rising watertable. duration of peat formation the original peat deposits in the øresund area were between 1.6 and 7.0 m thick and the peat-forming environments existed for between 1600 and 7000 years using the assumptions discussed above. although some clay sand sandsilt øresund-5 well upper part: open freshwater lake. limnic to limnotelmatic facies. mainly allochthonous deposition. lower part: mesotrophic to ombrotrophic domed bog. terrestrial facies. shallow freshwater lake freshwater lake m 52 55 56 57 58 m 45 47 48 49 50 53 54 59 clay silt 46 51 fig. 14. sedimentological core-log of the jurassic succession of the øresund-5 well. modified after petersen (1994); reference level: metres below sea level. for legend, see fig. 4. 649 of the mires were rather short-lived, the data demonstrate that certain of the middle jurassic peat mires in the øresund area were able to exist for longer periods than the early jurassic mires in the same area. discussion factors controlling peat formation peat formation occurred in the fennoscandian border zone from north jylland, through øresund and skåne to bornholm during early–middle jurassic times. the conditions for peat accumulation were most favourable to the south-east as indicated by an increase in number and thickness of the seams from north jylland to bornholm. the upper pliensbachian – lower toarcian sorthat formation coals differ from the bathonian upper bagå formation coals in terms of composition, thickness and palynological content. the lower jurassic coal seams represent peat accumulation in low-lying interchannel and abandoned channel mires on a lower coastal plain in an overall transgressive setting (petersen & nielsen 1995). in paralic settings, rising relative sea-level causes a rise in the groundwater table and a more landwards ponding of coarse sediments, which may result in the formation of extensive peatlands (diessel 1994; petersen & andsbjerg 1996; petersen et al. 1998). peat accumulation in the sorthat precursor mires was favoured by a continuously high-standing watertable and anoxic conditions, and only small amounts of very fine-grained detrital mineral particles reached the mires. the mires were probably rheotrophic and periodically saline-water influenced, and the resulting coals are generally very huminite-rich. the limited thickness of the seams (up to 0.57 m, but generally between 0.08 m and 0.26 m) reflects a rather short duration of peat accumulation (average for all seams of 2300 years), determined by a fairly rapid relative sea-level rise. the sea-level rise was governed by an overall eustatic rise (hallam 1988; haq et al. 1988) and a subsidence rate of c. 30 m/my (petersen clay sand sandsilt øresund-7 well lower to upper part: densely vegetated mesotrophic swamp. telmatic to terrestrial freshwater facies. lowermost part: densely vegetated rheotrophic infilled lake swamp. limnotelmatic to telmatic freshwater facies. freshwater lake densely vegetated mesotrophic swamp. telmatic freshwater facies. densely vegetated mesotrophic to ombrotrophic domed bog. (telmatic to) terrestrial freshwater facies. shallow freshwater lake densely vegetated rheotrophic infilled lake swamp. limnotelmatic to telmatic freshwater facies. shallow freshwater lake shallow freshwater lake possibly ombrotrophic raised bog. terrestrial facies. seam 1 seam 4 seam 3 seam 2 m 50 51 52 53 54 55 56 m 45 46 47 48 49 coal ? clay silt 44 fig. 15. sedimentological core-log of the jurassic succession of the øresund-7 well. modified after petersen (1994); reference level: metres below sea level. for legend, see fig. 4. & nielsen 1995). most of the coal seams represent the boundary between two succeeding lake/lagoonal successions, i.e. they were deposited during the transition from a period of relatively slow rise in the watertable to a period during which the watertable rose more rapidly. peat formation began when the lake/lagoon or abandoned channel was filled and turned into lowlying mires. peat accumulation probably continued during the initial phase of the succeeding faster rise in watertable, as accumulation of organic matter may have been favoured by sediment starvation due to upstream ponding of clastic sediments. peat formation was terminated by lacustrine or lagoonal flooding of the mires, when the watertable rise outpaced the rate of organic matter accumulation. similarly, the formation and duration of the hettangian–sinemurian coastal mires in the fennoscandian border zone was strongly controlled by a relative sea-level rise, and the peats obtained thicknesses of only 0.5–3.5 m. on average, the investigated mires in the rønne formation existed for c. 1200 years and the mires in the øresund area for 1850 years. the principal controlling factor on the development of the early jurassic mires in the fennoscandian border zone was thus a general rise of the groundwater table that favoured the development of waterlogged, possibly rheotrophic and nutrient-rich mires, and hindered the development of domed bogs in most cases. this resulted in the dominance of huminite-rich coals and carbonaceous claystones. a similar mechanism has been presented for the westphalian a in the warrior basin, alabama, where domed peat formation may have been hindered due to extremely rapid basin subsidence in spite of a suitably warm and humid climate (eble et al. 1994). hunt (1989) suggested that eastern australian permian delta plain coals are consistently high-vitrinite coals, because the peats formed under high watertables principally controlled by a high subsidence rate. during the middle jurassic, the subsidence rate was much lower as clearly shown by thinner middle jurassic sediment packages (up to 200 m in c. 21 ma; gradstein et al. 1994) in the danish basin and fennoscandian border zone compared to those in the lower jurassic (up to 1000 m in c. 26 ma). the marked change in subsidence rate and basin configuration was caused by uplift of the ringkøbing–fyn high and the danish basin south-west of the sorgenfrei–tornquist zone (nielsen 1995, 2003, this volume). these changes may have been related to thermal doming centrally in the north sea area (underhill & partington 1993) and to an important middle jurassic volcanic event in central skåne that was possibly associated with, or preceded by, uplift in the form of a rift dome (f. surlyk, personal communication 1997). the reduced subsidence rate coupled with a more stable eustatic sea level (hallam 1988; haq et al. 1988) resulted in a more inland position for the middle jurassic mires in the fennoscandian border zone and on bornholm than that for the early jurassic mires. the proximity of the bathonian mires (upper bagå formation) on bornholm to granitic basement is reflected in the coarser grain size of the quartz particles and a much higher kaolinite content than observed in the sorthat formation coals. the thickest of the upper bagå formation coal seams (1.9 m) corresponds to a precursor peat thickness of 19 m and an estimated 19 000 years of peat accumulation. from the øresund area, a maximum of 7000 years is estimated. this suggests that on occasion peat accumulation was able to keep pace with the increase in accommodation space for a relatively long period of time. at other times, the rate of peat accumulation may have outpaced the rate of watertable rise resulting in the development of domed peats. as noted above, this style of peat development probably resulted from the combination of a low subsidence rate and a stable eustatic sea level, which worked together to limit flooding controlled by relative sea-level rise. the coal petrography indicates that most of the investigated middle jurassic mires were exposed to fluctuations in the groundwater table, probably controlled by a seasonal climate; such periodic falls in the watertable level favoured oxidation of the peat surfaces. in contrast, the seasonal climate probably had a more limited influence on the level of the watertable in the early jurassic coastal mires, where a high-standing watertable was governed by a more or less continuous rise of relative sea level. oxidation of the organic matter on the middle jurassic peat surface implies that the estimates of peat accumulation periods may be a rough minimum. the combined effect of a seasonal climate and possibly peat accretion above the watertable level resulted in coal seams with a higher content of inertinite. a low subsidence rate has been similarly inferred to account for oxidation of the precursor peats of permian high-inertinite coals in eastern australia (hunt 1989; hunt & smyth 1989). the fluctuating watertable in the middle jurassic mires may have favoured wildfire frequency and the formation of fire-derived inertinite; however, the nature of the inertinite, particularly in the coals from the øresund area, is indicative of derivation by oxidation. the presence of inertinitic maceral precursors formed by oxidation in recent peat deposits (cohen & stack 1996; moore et al. 1996), makes this suggestion likely. 650 the peat-forming vegetation based on palynological evidence the palynological evidence of the peat-forming vegetation suggests a lower jurassic flora with a prominent proportion of ferns together with cycadophytes, coniferophytes from the families pinaceae and taxodiaceae, unknown gymnospermous plants, and bryophytes. although the overall composition of the plant communities seems to have been rather uniform, variations in the composition of the palynomorph assemblages (and coal composition) in the seams and between seams may reflect variations in the peat-forming plant communities. these variations are probably related to specific environmental conditions in the mires as demonstrated in recent peat systems (cohen 1973; esterle & ferm 1994; phillips & bustin 1996b). the limited knowledge of the ecology and environmental tolerance of the plants restricts the understanding of the palaeoecological differences between the mire systems as indicated by various palynomorph assemblages. the middle jurassic seams from the upper bagå formation are characterised by the common occurrence of spores whose affinities lie with the order lycopodiales and the absence of pollen related to the order cycadales. club mosses are particularly abundant in mineral-rich levels corresponding to wet, nutrient-rich conditions in the precursor mires, or common inundations. bryophytes seem to have been present, together with ferns and coniferophytes. of particular note is the absence of pollen related to cycadales in the middle jurassic coals, particularly in view of their common occurrence in the lower jurassic samples. according to lapo & drozdova (1989), cycadophytes are rarely found in jurassic coals because they preferred drier habitats. however, höhne (1933) noted the presence of cycadophytes in the bagå formation coals, although this may have been stratigraphically lower in the succession. it is notable that the genus nilssonia, considered to represent the foliage of cycadophytes, occurs in lower jurassic claystones on bornholm (möller 1903). hence the palynomorphs from the lower jurassic seams whose affinities are with the cycadales could have originated from vegetation outside the mires. in general, there is a relatively good correspondence between the interpretation of the peat-forming vegetation based on macerals and the palynological evidence of the peat-forming vegetation. in order to strengthen the interpretations based on macerals, however, it is necessary to acquire specific knowledge of the peat-forming plants, i.e. the stature, structure and the wood content of the tissues; furthermore, the degradational conditions should also be considered (pierce et al. 1995). biological aspects of the parent plants should be incorporated, if possible, in particular to account for the variation in spore/pollen production by different species, but also to account for spores and pollen that may have been transported into the mires from the hinterland. it is clear that multidisciplinary studies of coals integrating palaeobotanical data with petrographic results provide the most detailed picture of the mire floras (see discussion by scott 1991; dimichele & phillips 1994). origin and environmental significance of pyrite in the coals of the sorthat and bagå formations the framboidal pyrite in the sorthat formation coals is interpreted to be syngenetic. framboidal pyrite is normally considered to indicate the influence of saline water during peat formation because organic sulphur inherited from the peat-forming flora only accounts for low contents of sulphur and pyrite in freshwater environments (cohen et al. 1984; casagrande 1987; phillips & bustin 1996a). however, the nature of the pyrite in the bagå formation coals and the lack of other evidence for marine influence suggest that this pyrite is mainly epigenetic in origin. it is therefore necessary to identify an external source of sulphur to account for the high content of pyrite in fairly well-defined intervals in the coals and the abundant pyrite concretions in the sand beds. the coal-bearing strata are faulted against the graben bounding fault, and sulphate-enriched porewater expelled from deeper levels may have migrated up through the fault zone and into the highly permeable sand beds and coal seams. sulphate-rich porewater is associated with anhydrite deposits, and zechstein and triassic deposits with anhydrite and a very low content of organic matter are present in the rønne graben. sulphate-rich porewater is known from triassic strata in the danish basin (gustafson & anderson 1979; laier 1982). the pyrite nodules in the sand beds in the bagå beds are often associated with coal particles. this suggests that the reduction of the sulphate and the precipitation of pyrite occurred where the local environment was reducing due to enrichment in organic matter. in the coal seams, pyrite is often found associated with fusinite, suggesting that precipitation was favoured by the higher porosity of the fusinite compared to the huminite (fig. 8d). an analogous situation has been described from carboniferous coal seams in a non651 marine succession from nova scotia, canada, where epigenetic pyrite fills fusain lenses, fractures and cleats (beaton et al. 1993). the carboniferous coal measures are in fault contact with rocks containing gypsum, anhydrite, carbonates, halite and potash salts, which are interpreted to be the source of the sulphur. conclusions the development of the early–middle jurassic mires in the fennoscandian border zone occurred in a humid, warm-temperate to subtropical, weakly seasonal climate. during the early jurassic, mires were established on coastal plains facing the open sea to the west. peat formation was mainly controlled by an overall fairly rapid rise of relative sea level governed by the general eustatic rise and a relatively high subsidence rate causing repeated lacustrine or lagoonal flooding of the mires. the early jurassic mires were relatively short-lived, averaging 1200 years (rønne formation), 1850 years (øresund area) and 2300 years (sorthat formation), because watertable rise outpaced the rate of peat accumulation. the lower jurassic coals and coaly mudstones in the fennoscandian border zone and on bornholm represent peat deposits that were between 0.5 and 5.7 m thick, but generally were less than 3 m thick. the upper pliensbachian – lower toarcian sorthat formation coals originated in continuously waterlogged, anoxic mires, and the high-standing watertable and the periodic marine influence on the mires favoured the formation of huminite-rich coals and the precipitation of framboidal pyrite. in addition to pyrite, the mineral matter is dominated by clay and fine-grained silt-sized quartz. uplift of the ringkøbing–fyn high and the main part of the danish basin south of the sorgenfrei–tornquist zone in aalenian times changed the basin configuration and influenced the accumulation of peat. the uplift, which may have been related to the formation of a large thermal dome in the central north sea, and possibly also in skåne, caused a strongly reduced subsidence rate within the fennoscandian border zone. the low subsidence rate coupled with a relative stable eustatic sea level, resulted in a general regression such that the middle jurassic mires developed farther inland than those of the early jurassic. the setting enabled peat accumulation in the middle jurassic mires to keep pace with and occasionally outpace the watertable rise. thus, some of the middle jurassic peat mires existed for relatively long periods. it is inferred from coals in the øresund area that peat accumulation occurred for up to 7000 years and upper bagå formation coals testify that at certain times during the bathonian, peat accumulation may have occurred on bornholm without significant interruption for up to 19 000 years. the combination of an inland position of the middle jurassic mires, a seasonal climate, and periodic peat accretion above the groundwater level caused temporary oxidation of the peat surfaces and the formation of inertiniterich coals. watertable fluctuations within the upper bagå formation mires on bornholm may further have been promoted by the generally sandy substrate of the alluvial plain. the proximity of these mires to the granitic basement areas to the east on bornholm is reflected by a larger grain size of quartz particles and a higher kaolinite content than observed in the sorthat formation coals. the pyrite in the upper bagå formation coals is interpreted to be mainly epigenetic in origin, precipitated out of sulphate-enriched porewater expelled from underlying zechstein–triassic strata. palynological data indicate variations in the peatforming plant communities between coal seams. variations within individual seams also suggest subtle changes in the mire environment, probably mainly of hydrological character, reflecting the evolution of the mire. the hettangian–sinemurian mires represented by the seams in the øresund-18 well were characterised by an abundance of ferns, cycadophytes and at certain levels significant proportions of coniferophytes of the family taxodiaceae. club mosses were also represented. the late pliensbachian – early toarcian mires (seams 2 and 15 in the levka-1 well, sorthat formation) were also characterised by an abundance of ferns. in contrast to the sinemurian mires in the øresund area, however, cycadophytes were more abundant, while coniferophytes of the family taxodiaceae were much less common and club mosses were probably absent. the bathonian mires represented by the seams in the upper bagå formation differ in particular from the early jurassic mires by the complete absence of cycadophytes and the abundance of club mosses, and by the rare occurrence of coniferophytes of the family taxodiaceae. ferns were significant components of the mires, as were jurassic species of pinaceae. it is tempting to relate the difference in the plant communities of the mires represented by the sorthat and bagå formation coal seams to the distinctly different depositional settings represented by these seams. the early jurassic mires were waterlogged, anoxic and periodically marine-influenced, whereas the bathonian mires developed under fresh652 water conditions and were situated inland, where in particular the seasonal climate was important in promoting a fluctuating watertable. in general, there appears to be a relationship between the petrographic composition of the coals and the composition of the peatforming vegetation, as deduced from palynological evidence. thus, the overall coal seam composition was controlled not only by allogenic factors but also by differences in vegetational input from the various peat-forming plant communities, which may account for the more subtle petrographic variations in the coal seams. acknowledgement the paper benefited significantly from constructive reviews by d.j. 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(2015). we estimate ice-sheet output, or the ice discharge across the ice-sheet grounding line, by applying downstream corrections to the ice flux across the promice perimeter. we subtract this ice discharge from ice-sheet input, or the area-integrated, ice sheet surface mass balance, estimated by a regional climate greenland ice sheet mass balance assessed by promice (1995–2015) william colgan*1, kenneth d. mankoff1, kristian k. kjeldsen1,2, anders a. bjørk2, jason e. box1, sebastian b. simonsen3, louise s. sørensen3, s. abbas khan3, anne m. solgaard1, rene forsberg3, henriette skourup3, lars stenseng4, steen s. kristensen5, sine m. hvidegaard3, michele citterio1, nanna karlsson1, xavier fettweis6, andreas p. ahlstrøm1, signe b. andersen1, dirk van as1 and robert s. fausto1 geus bulletin is free to individuals and institutions in electronic form. the author(s) retain copyright over the article contents. research article | open access geus bulletin vol 43 | e2019430201 | published online: 08 july 2019 https://doi.org/10.34194/geusb-201943-02-01 fig. 1. a the promice perimeter and nineteen ice sheet sub-sectors (zwally et al. 2012). major glaciers (jakobshavn (jak), humboldt (hum), zachariae (zac), kangerlussuaq (kan) and helheim (hel)) are shown for reference. b average surface mass balance (smb) in mm of water equivalent (we) per year over the 1980–1999 period from four mar3.5.2 simulations (fettweis et al. 2017). c satellite-derived ice-surface velocity during winter 2008/2009 (rignot & mouginot 2012). both surface mass balance and ice-surface velocity data are only shown within the promice ice-sheet mask, excluding independent ice caps and glaciers (citterio & ahlstrøm 2013). 300 km 50ow 75on 70on 65on 60on 40ow 70ow 50ow 30ow 10ow 0o 1000 100 10 velocity (m/yr) 10ow30ow 40ow50ow 50ow70ow 0o 300 km smb (mmwe/yr) 2000 –2000 0 70ow 50ow 30ow 10ow 0o a b c 300 km 50ow 80on 75on 70on 65on 60on 40ow 1.1 1.2 1.3 1.4 2.1 2.2 3.1 3.23.3 4.1 4.2 4.3 5 6.1 6.2 7.1 7.2 8.1 8.2 hum zac jak kan hel km 0 km 5415 https://doi.org/10.34194/geusb-201943-02-01 e2019430201-02 model. while andersen et al. (2015) assessed ice-sheet mass balance in 2007 and 2011, this updated input-output assessment now estimates the annual sea-level rise contribution from eighteen sub-sectors of the greenland ice sheet over the 1995–2015 period. input-output method ice discharge is calculated as the ice flux across the promice perimeter, corrected for downstream mass changes due to surface mass balance and changing ice volume (fig. 2). ice flux (f) across the promice perimeter within a given ice-sheet sub-sector is calculated as: f = ∑l i=0 vpi hi ρf ∙ ∆l where vp is gate-perpendicular, or perimeter-perpendicular, ice-surface velocity, h is the ice thickness, ρ is bulk ice-sheet density (assumed to be 915 ± 2 kg/m3) and f is the ratio of surface-to-depth-averaged ice velocity (assumed to be 0.93 ± 0.05; thomas et al. 2001). ice flux is summed along the promice perimeter length (l) within a given ice-sheet sub-sector in increments (δl) of 30 m. gate-perpendicular velocity is calculated as: vp= v cos ϑ where v is the absolute surface velocity and ϑ is the difference between ice flow and gate-perpendicular azimuths. when ϑ exceeds 90°, gate-perpendicular velocity becomes negative, indicating ice flow into the perimeter (fig. 3). this reverse ice inflow occurs along 5.1% of the entire promice perimeter (275 km), primarily in east greenland. grounding-line ice discharge (d) is calculated as the sum of ice flux across the promice perimeter (f) and two downstream corrections that account for changing ice volume and surface mass balance: d = f – v . ds + b . ds where v . ds is the area-integrated observed rate of change in ice volume downstream of the perimeter, and b . ds is the areaintegrated modelled surface mass balance downstream of the perimeter. the rate of change in downstream ice volume captures changes due to both surface mass balance and ice dynamics. this requires the secondary surface mass balance correction to isolate the ice dynamic contribution to grounding-line ice discharge. subtracting a negative downstream volume change increases ice flux, while adding a negative downstream surface mass balance decreases ice flux. we assess mass balance (m. ) within a given ice-sheet subsector as: m. = b . – d where b . is area-integrated modelled surface mass balance and d is calculated grounding-line ice discharge. we assess mass balance in eighteen of nineteen ice-sheet sub-sectors delineated by zwally et al. (2012) using the promice ice-sheet mask (citterio & ahlstrøm 2013). these eighteen minor sub-sectors are aggregated into eight major sectors (fig. 1a). we do not assess mass balance in sector 3.2 (geikie plateau). we propagate the uncertainties following andersen et al. (2015), whereby we employ quadratic sums for terms with common units and quadratic fractional sums for terms with differing units. datasets we interpolate satellite-derived synthetic aperture radar ice velocity (v) along the promice perimeter from a spatially complete and temporally constrained winter 2008/2009 velocity mosaic of 150 m spatial resolution (rignot & mouginot 2012). where possible – along 67% of the promice perimeter – we derive temporal trends in ice velocity from overlapping winter 2008/2009 and winter 2014/2015 velocity mosaics (joughin et al. 2010). we apply these temporal trends within each sub-sector to estimate annual perimeter velocity profiles during the 2000–2015 period. this approach is meant to complement the spatial completeness of the rignot & mouginot (2012) annual mosaic with the temporal repeat of the joughin et al. (2010) data. we assume the perimeter velocity profile in the year 2000 is characteristic of the 1995–1999 period, on the basis that the ice-sheet interior was near equilibrium mass balance prior to 2000 (thomas fig. 2. schematic representation of calculating ice discharge using the promice perimeter. ice discharge (d) at the grounding line is derived from downstream rate of change in ice volume (v .ds ) and surface mass balance (b . ds) corrections applied to the ice flux (f) observed through the promice perimeter (andersen et al. 2015). e2019430201-03 et al. 2001). these simplifications overlook pre-2008 ice flow variability, such as the acceleration and deceleration of south-east greenland glaciers during 2000–2007 (enderlin et al. 2014). while pre-2008 ice-sheet velocity maps are available, their quality decreases inland from the ice-sheet margin, which results in poor sampling along the promice perimeter (fig. 3). we estimate ice thickness (h) along the promice perimeter using ice surface and bed elevation data. where possible, ice thickness is calculated from promice airborne campaign laser and radar altimetry measurements in 2007, 2011 and 2015 (sørensen et al. 2018). promice airborne radar surveys have measured bedrock elevation along 79% of the perimeter. bedrock elevations are interpolated along the remaining 21% of the perimeter from bedmachine v3 (morlighem et al. 2017). promice laser altimetry surveys each measured ice-sheet surface elevation along 74 to 79% of the perimeter. along 15% of the perimeter never surveyed by airborne altimetry, ice-sheet surface elevations are interpolated from a digital elevation model representative of 2007 (howat et al. 2014). when and where required, independent altimetry-derived rates of elevation change are used to derive annual elevation profiles along the perimeter during 2000–2015 (khan et al. 2016). we interpolate area-integrated rate of change in ice volume observed downstream of the promice perimeter (v . ds) annually within each ice-sheet sub-sector during the 1995–2015 period from the same independent airand satellite-borne altimetry product (khan et al. 2016). these rates of volume change have been corrected for firn compaction when and where necessary. we estimate rates of change in ice volume in each sub-sector by area-integrating this altimetry product, and associated uncertainties, at 500 m spatial resolution, over the ice-sheet area downstream of the promice perimeter (citterio & ahlstrøm 2013). fig. 3. a: satellite-derived annual winter icevelocity data availability along the promice perimeter within the combined joughin et al. (2010) and rignot & mouginot (2012) datasets during winters 1995–2015. b: 2008/2009 ice flow and gate-perpendicular azimuths around the promice perimeter (rignot & mouginot 2012). c: dimensionless scale factor (cos ϑ) of velocity magnitude. d: absolute and gate-perpendicular surface velocity around the promice perimeter. e: changes in gate-perpendicular ice velocities surveyed in 2008/2009 and 2014/2015 (joughin et al. 2010). in all subplots vertical dashed lines denote major glaciers (fig. 1). e2019430201-04 we use surface mass balance simulated by the mar3.5.2 regional climate model for both downstream surface mass balance correction (b . ds) and assessing ice-sheet wide surface mass balance input (b . ). this permits us to assimilate the runoff and snowfall rates of a four-simulation ensemble reflecting four different climate forcings (era-20c, era-interim, ncepv1 and 20crv2c) into an annual surface mass balance time series that spans 1980–2015 at 500 m spatial resolution. we remove relative anomalies between these four simulations during the common 1980–1999 period (fettweis et al. 2017). the promice ice-sheet mask we employ has a more extensive ice-sheet ablation area than the native mar3.5.2 ice mask (citterio & ahlstrøm 2013). relative to the native 25 km mar ice mask, the more extensive ice-sheet ablation area of the 500 m promice ice mask decreases ice-sheet integrated surface mass balance by c. 30 gt/yr. the ice-sheet integrated downscaled surface mass balance we interpolate is within the range of independently elevationdependent downscaled mar2 simulations (franco et al. 2012). ice sheet mass loss our updated input-output assessment gives a total 1995– 2015 ice-sheet mass loss of 3028 ± 711 gt (fig. 4). this is equivalent to a eustatic sea-level rise contribution of 8.4 ± 1.9 mm. we assess all eight major ice-sheet sectors as within uncertainty of equilibrium balance at the start of the surfig. 4. annual surface mass balance, ice discharge and mass balance in eight major ice-sheet sectors (1–8), as well as for the entire ice sheet, over the 1995– 2015 period. vertical spread denotes associated uncertainty. map: spatial distribution of the average 1995–2015 mass balance (mb) within the promice ice-sheet mask (citterio & ahlstrøm 2013; khan et al. 2016). black lines denote the eight major ice-sheet sectors (zwally et al. 2012). 1995 2000 2005 2010 2015 (g t/y r) -100 0 100 1 1995 2000 2005 2010 2015 (g t/y r) -100 0 100 2 1995 2000 2005 2010 2015 (g t/y r) -100 0 100 3 1995 2000 2005 2010 2015 (g t/y r) -100 0 100 4 1995 2000 2005 2010 2015 (g t/y r) -100 0 100 5 1995 2000 2005 2010 2015 (g t/y r) -100 0 100 6 1995 2000 2005 2010 2015 (g t/y r) -100 0 100 7 1995 2000 2005 2010 2015 (g t/y r) -100 0 100 8 1995 2000 2005 2010 2015 (g t/y r) -600 -400 -200 0 200 400 600 ice sheet ice discharge surface mass balance mass balance 0°10°w30°w 40°w50°w 50°w70°w 80°n 75°n 75°n 70°n 70°n 65°n 65°n 60°n 60°n300 km mm/yr mb 2000 0 -2000 2 3 7 6 4 5 1 8 e2019430201-05 vey period (c. 1995). negative mass balance years, however, have clearly become more common towards the end of the survey period (fig. 4). in particular, marine-terminating sectors with substantial ice discharge (central west (7), southeast (4) and north-east (8)) transitioned to persistent mass loss c. 2002, 2004 and 2005, respectively. land-terminating sectors with substantial meltwater runoff (south (5), southwest (6)) subsequently transitioned to persistent mass loss c. 2006. at the ice-sheet scale, the total mass loss we assess over the 1995–2011 period agrees, within uncertainty, with that assessed by the ice sheet mass balance inter-comparison exercise (imbie) (fig. 5; shepherd et al. 2012). the apparent discrepancy between imbie and promice mass loss estimates is approximately equivalent to independent estimates of peripheral glacier mass loss (noël et al. 2017). adding the promice mass loss estimate for the ice sheet proper with the peripheral glacier mass loss estimate of noël et al. (2017) suggests that peripheral glaciers were responsible for 17 ± 7% of greenland’s contribution to sea level change during 2004–2013. peripheral glaciers account for < 5% of greenland’s ice-covered area (citterio & ahlstrøm, 2013), making their specific, or per unit area, sea-level contribution disproportionately greater than the ice sheet. our linear extrapolation of pre-2008 ice velocities, under-sampling flow variations in south-east greenland glaciers in 2000–2007, likely contributes to some discrepancy with the imbie sealevel contribution curve (enderlin et al. 2014). ice discharge increased from 1995 (350 ± 72 gt/yr) to 2009 (487 ± 71 gt/yr), before decreasing slightly to 2015 (465 ± 74 gt/yr). persistently increasing trends in iceberg calving are more readily apparent in sectors 7 and 8 than in sectors 3 (central east) and 4 (fig. 4). the trend in ice flux across the promice perimeter ( 1.5 gt/yr/yr) was small in comparison to the trend in ice discharge across the grounding line ( 7.2 gt/yr/yr) during 2000–2015. the majority of the inter-annual variability in grounding-line ice discharge therefore results from the downstream surface mass balance and ice volume corrections we apply to the perimeter flux. ice-sheet wide ice discharge during 2000–2015 (432 ± 74 gt/yr) is c. 45 gt/yr (10%) lower than that assessed by king et al. (2018) (479 ± 20 gt/yr) for the same period. during 2000–2010, our ice discharge (422 ± 74 gt/yr) is c. 90 gt/yr (21%) lower than that assessed by enderlin et al. (2014; 511 ± 30 gt/yr) during the same period. the formal uncertainty of a given study can therefore be substantially smaller than inter-study discrepancies. trends and variability in mar-simulated ice-sheet wide surface mass balance have been widely discussed (fettweis et al. 2017). the ice-sheet wide annual surface mass balances we employ are consistent with the relatively extensive ablation area of the 500 m resolution promice ice mask and the average surface mass balance we interpolate over the 1990–2010 period (382 ± 58 gt/yr) is within the sensitivity range of the elevation-dependent downscaled product of mar2 simulations from their native 25 km resolution to 15 km resolution during the same period (franco et al. 2012). as virtually the entire ice-sheet ablation area is downstream of the promice perimeter, the ice discharge we assess is fundamentally dependent on simulated surface mass balance. a more positive surface mass balance simulation would result in greater ice discharge and vice versa. differences in downstream surface mass balance correction are primarily responsible for the c. 55 gt/yr (12%) decrease in the ice discharge assessed here (460 ± 75 gt/yr) in comparison to that originally assessed by andersen et al. (2015; 515 ± 57 gt/yr) during 2007 and 2011. programme outlook this report updates the contribution of the greenland ice sheet to annual sea-level rise assessed by promice, using the andersen et al. (2015) input-output approach. we assess an ice-sheet mass loss of 3028 ± 711 gt over the 1995–2015 period, which is equivalent to a eustatic sea-level rise contribution of 8.4 ± 1.9 mm. combining our estimate of ice-sheet mass loss with a previous estimate of peripheral glacier mass loss yields a total greenland ice loss (ice sheet plus peripheral fig. 5. cumulative sea-level equivalent (sle) contribution from this promice study shown in comparison to the greenland mass balance inter-comparison exercise (imbie; shepherd et al. 2012). as imbie surveys both the ice sheet and peripheral glaciers, we also sum this study with the independent peripheral glacier contribution estimate of noël et al. (2017) for context. 1995 2000 2005 2010 2015 se ale ve l c on tr ib ut io n (m m ) 0 2 4 6 8 10 12 imbie promice (this study) promice + noël2017 e2019430201-06 *corresponding author: william colgan | e-mail: wic@geus.dk 1 department of glaciology and climate, geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350, copenhagen k, denmark. 2 natural history museum, university of copenhagen, copenhagen, denmark 3 department of geodynamics, technical university of denmark, lyngby, denmark 4 department of geodesy, technical university of denmark, lyngby, denmark 5 department of microwave and remote sensing, technical university of denmark, lyngby, denmark 6 department of geography, university of liège, liège, belgium glaciers) that is consistent with the most recent consensus of total greenland ice loss (shepherd et al. 2012; noël et al. 2017). digital versions of the area-integrated calendar year mass balance that we assess in eighteen ice-sheet sub-sectors, as well as underlying components, are available on the www. promice.dk website. as a result of the relatively high-elevation and inland location of the promice perimeter, virtually the entire icesheet ablation area resides downstream. this makes interannual variability in grounding-line ice discharge estimated by andersen et al. (2015), highly sensitive to inter-annual variability in downstream corrections. future promice mass balance products will therefore adopt new approaches where ice flux is estimated across gates near the grounding lines of individual outlet glaciers. the sustained effort of the programme for monitoring of the greenland ice sheet (promice) will continue to provide danish and international stakeholders open access to policy-relevant estimates of ice-sheet mass loss and sea-level rise. acknowledgements this work is a product of the programme for monitoring of the greenland ice sheet (www.promice.dk), which is funded by the danish cooperation for environment in the arctic (dancea) through the danish ministry of climate, energy and utilities. we thank the reviewers, rachel carr and ellyn enderlin, for their comments, which improved the manuscript. references andersen et al. 2015: basin-scale partitioning of greenland ice sheet mass balance components (2007–2011). earth and planetary science 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higher-elevation parts of the greenland ice sheet. journal of geophysical research 106, 33,707–33,716. https://doi.org/10.1029/2001jd900033 zwally, h., giovinetto, m., beckley, m., & saba, j. 2012: antarctic and greenland drainage systems. digital media. http://icesat4.gsfc.nasa. gov/cryo_data/ant_grn_drainage_systems.php how to cite colgan, w., mankoff, k.d., kjeldsen, k.k., bjørk, a.a., box, j.e., simonsen, s.b., sørensen, l.s., khan, s.a., solgaard, a.m., forsberg, r., skourup, h., stenseng, l., kristensen, l.s., hvidegaard, s.m., citterio, m., karlsson, n., fettweis, x., ahlstrøm, a.p., andersen, s.b., van as, d., fausto1, r.s. 2019: greenland ice sheet mass balance assessed by promice (1995–2015). geological survey of denmark and greenland bulletin 43, e2019430201. https://doi.org/10.34194/geusb-201943-02-01 mailto:wic%40geus.dk?subject= https://doi.org/10.22008/promice/data/mass_balance/input_output/perimeter 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http://icesat4.gsfc.nasa.gov/cryo_data/ant_grn_drainage_systems.php https://doi.org/10.34194/geusb-201943-02-01 geological survey of denmark and greenland bulletin 1, 21-59 stratigraphy previous page: lower jurassic ammonites (platypleuroceras caprarium) from bornholm, denmark – see donovan & surlyk (2003, this volume). photo: kristian kloth-jørgensen. 23 the lower jurassic of europe: its subdivision and correlation kevin n. page the lower jurassic sub-system comprises four stages, in chronological order, the hettangian, sinemurian, pliensbachian and toarcian. each stage is subdivided into a sequence of ‘standard zones’ (= chronozones) and subzones – each correlated primarily on the basis of its ammonite fauna. a further increase in stratigraphical resolution is available by the use of intra-subzonal units known collectively as ‘horizons’. the close link between ammonites and chronostratigraphy means that faunal provincialism may determine which zonal framework, and therefore which subdivision of the lower jurassic, applies in different regions of europe. such provincialism is of minor importance in the early jurassic (hettangian – lower pliensbachian) but increases significantly in the upper pliensbachian and into the toarcian where at least three ammonoid faunal provinces are distinguishable. the standard zonal schemes for each relevant faunal area are discussed here, with greatest emphasis being placed on the northwest european province, which is characteristic of much of northern europe throughout most of the early jurassic. intra-subzonal units have only been described in certain regions for parts of the lower jurassic but where recognisable these are introduced. keywords: europe, lower jurassic, ammonite zonal biostratigraphy department of geological sciences, university of plymouth, drake circus, plymouth, devon pl4 8aa, uk. e-mail: kevinp@bello-page.fsnet.co.uk contents d’orbigny’s stages, oppel zones and the lower jurassic chronostratigraphy: stages, standard zones, subzones, chronozones and ammonites intra-subzonal units: biohorizons and zonules ammonoid provincialism in the lower jurassic of europe: consequences for correlation hettangian stage the base of the hettangian stage and the jurassic system hettangian correlative schemes and ammonoid provincialism northwest european province (britain, ireland, france, germany, etc.) planorbis chronozone liassicus chronozone angulata chronozone mediterranean province (austria) calliphyllum chronozone megastoma chronozone marmorea chronozone 25 25 26 26 27 27 28 29 29 30 30 31 31 31 31 geological survey of denmark and greenland bulletin 1, 23–59 (2003) © geus, 2003 sinemurian stage the base of the sinemurian stage sinemurian correlative schemes and ammonoid provincialism northwest european province (britain, ireland, france, germany, switzerland, etc.) lower sinemurian substage bucklandi chronozone semicostatum chronozone turneri chronozone upper sinemurian substage obtusum chronozone oxynotum chronozone raricostatum chronozone mediterranean province (austria, italy) pliensbachian stage the base of the pliensbachian stage pliensbachian correlative schemes and ammonoid provincialism northwest european province (britain, france, germany, northern spain, etc.) lower pliensbachian substage jamesoni chronozone ibex chronozone davoei chronozone northwest european province: lusitanian ?sub-province (portugal) subboreal province (britain, northern germany, etc.) upper pliensbachian substage margaritatus chronozone spinatum chronozone submediterranean province (southern france, northern spain, etc.) mediterranean province (austria, italy, southern spain) toarcian stage the base of the toarcian stage toarcian correlative schemes and ammonoid provincialism subboreal province (northern britain) lower toarcian substage tenuicostatum chronozone serpentinum chronozone bifrons chronozone submediterranean province (southern england, france, germany, northern spain, etc.) lower toarcian substage tenuicostatum chronozone serpentinum chronozone bifrons chronozone northwest european province (britain, france, germany, northern spain, etc.) upper toarcian substage variabilis chronozone thouarsense chronozone dispansum chronozone pseudoradiosa chronozone aalensis chronozone mediterranean province (italy, austria, southern spain, north africa, etc.) lower toarcian substage polymorphum chronozone levisoni chronozone bifrons chronozone upper toarcian substage gradata chronozone bonarellii chronozone speciosum chronozone meneghini chronozone aalensis chronozone references 24 31 31 32 32 32 32 34 34 34 34 36 36 37 38 38 38 39 39 39 39 41 41 41 41 41 43 43 43 43 44 44 45 45 45 46 46 47 47 47 48 48 49 49 49 50 50 51 51 51 51 51 52 52 52 52 52 52 53 53 53 the jurassic system takes its name from the jura mountains of eastern france and switzerland, first recognised as having rocks of a distinctive geological division by alexander von humbolt in 1795. it was not until 1829, however, that the term ‘jurassique’ was introduced by brongniart (1829), although in a restricted sense essentially for what is now known as the middle jurassic (torrens & getty 1980). inclusion of what is now lower and upper jurassic was soon made and by 1842, with the publication of alcide d’orbigny’s ‘palaeontologie française, terrains jurassique’ (1842–1849), a system of subdivision of a modern-looking jurassic system into stages was well under way (arkell 1933; rioult 1974). d’orbigny’s stages, oppel zones and the lower jurassic the stages introduced by d’orbigny (1842–1849) were designed to be of worldwide use and were based on the assumption that periodic mass extinctions followed by rapid re-establishments of new faunas characterised stage boundaries (arkell 1933, p. 9). the duration of the existence of a particular ‘fauna’ therefore correlated rocks belonging to a specific stage. for what is now considered to be lower jurassic, he recognised three successive stages: sinémurien (after semur-en-auxois in burgundy, eastern france), liasien (derived from the old lithological or lithostratigraphical term ‘lias’) and toarcien (after thouars in western france). d’orbigny (1842–1849) included in each of his stages a series of fossil ‘zones’, using the term in a purely palaeontological sense to indicate the general stratigraphic range of particular taxa. the term ‘zone’ was refined by albert oppel (1856–1858), who developed a sequence of such divisions for the entire jurassic system. oppel also apparently firmly established the use of lower jurassic (‘unterer jura’) in a modern sense, as being equivalent to the earlier essentially lithostratigraphical division known as ‘lias’. although not specifically using d’orbigny’s stages, oppel referred to zonengruppen or etagen. for the lower jurassic he recognised three successive divisions of this type: semur-gruppe (equivalent to d’orbigny’s sinémurian), pliensbach-gruppe (equivalent to liasien, named after pliensbach in württemberg, germany) and thouars-gruppe (equivalent to toarcian). very importantly, oppel established the ‘zone des ammonites planorbis’ as marking the base of the jurassic. this convention still stands, as does oppel’s general framework of zones, the only significant subsequent change being the creation of the hettangian stage by renevier (1864) for the first two zones of oppel’s original scheme. many other stage names have been proposed for parts of the lower jurassic in europe, some have continued to be used for divisions at the level of substage, but most are now redundant. a full list of such terms was provided by arkell (1933), together with extensive discussion of the origins of the various schemes for subdividing the jurassic system. conventionally the sinemurian, pliensbachian and toarcian stages are divided into an upper and a lower substage (dean et al. 1961), generally used without a specific identifying name. nevertheless, substage names are occasionally used and these are introduced in the relevant sections below. chronostratigraphy: stages, standard zones, subzones, chronozones and ammonites chronostratigraphy is the establishment of a series of subdivisions of geological time, using actual rock units as standards for reference. there is much discussion of this method and it is not necessary to develop it further here (see, for example, hedberg 1976; callomon 1984; harland et al. 1990; salvador 1994; remane et al. 1996). chronostratigraphic divisions are defined only at their base in a suitable stratotype section, the top of the unit being identified by the actual or correlated base of the next equivalent ranked division of the scale. chronostratigraphical divisions form a hierarchy with systems, series and stages being three divisions of decreasing rank (although the term series is almost never used in jurassic stratigraphy). the definitions of stages and systems is now regulated by the international commission on stratigraphy (ics), a project of iugs/unesco, through subcommissions focused on single systems. the aim is to formally recognise an internationally agreed global stratotype section and point (gssp) for the base of every system and for every stage of every system (cowie et al. 1986; salvador 1994). several such proposals have now been ratified by iugs, including the sinemurian stage of the lower jurassic, and several others are likely to be agreed within the next few years. below the level of stage, subdivisions at the level of chronozone and ultimately zonule can be used, but are not formally regulated through the international subcommission on jurassic stratigraphy (isjs) or the ics. in the jurassic, the often great abundance of ammonites and their wide geographical distribution has led to their 25 use for correlating sequences of standard zones. as discussed at great length elsewhere (callomon 1965, 1984; callomon & donovan 1974; cox 1990), these standard zones are chronozones and should therefore be treated as such – a fact ignored by some authors (e.g. whittaker et al. 1991) who confuse jurassic ammonite zones with biozones, where the use of fossils in correlation is not explicitly linked to geological time. as discussed by the former authors, although the names of the zonal units are derived from species names, they are by convention quoted non-italicised (e.g. jamesoni chronozone or jamesoni [standard] zone and not uptonia jamesoni zone or biozone). this is the convention followed by working groups of the international subcommission on jurassic stratigraphy. other fossil groups, especially microfossil, have been used to construct true biozonal schemes for the marine jurassic but the resolution of these schemes is usually inferior to the ammonite scale. indeed, the latter scale is typically used as a ‘standard’ against which biozonal schemes are correlated. for this reason, only the ammonite-based standard zonations for europe will be considered further here; microfossil schemes for the jurassic of northwest europe were reviewed by cox (1990) and dommergues (1997), based on dinocysts (woollam & riding 1983; falconnier 1997), calcareous nannofossils (bown et al. 1988; gardin 1997), foraminifera (copestake & johnson 1989; bassoullet 1997; ruget & nicollin 1997) and ostracods (bate & coleman 1975; lord 1978; park 1984; bodergat 1997; colin 1997). tables 71 and 72 in dommergues (1997) provide a recent cross-correlation between these various schemes, and others based on macrofossil groups – the latter including belemnites (based on doyle 1990; combémorel 1997), brachipods (based on alméras et al. 1997) and echinoderms (based on thierry et al. 1997). intra-subzonal units: biohorizons and zonules most jurassic chronozones are divided into subchronozones, largely for historical reasons, as the creation of new subchronozones within existing chronozones achieves a degree of nomenclatural stability at a chronozonal level. smaller divisions than subchronozones, however, are also used in jurassic ammonite stratigraphy and their use to further refine correlations again avoids ‘tampering’ with an established standard zonation. such divisions are generally known collectively (and sometimes confusingly) as ‘horizons’ although including two conceptually different types of unit (page 1995a). the first type of horizon, known as a zonule (as adopted by phelps (1985) following hedberg (1976), is the smallest subdivision of a chronostratigraphical scale. it should therefore be defined, as with higher divisions, by a basal boundary stratotype. the second type of unit is a biohorizon and is defined as “a bed or series of beds, characterised by a fossil assemblage, within which no further stratigraphical differentiation of the fauna or flora can be distinguished” (callomon 1984, p. 624). the earlier term hemera, proposed by buckman (1893), is considered to be the chronological equivalent of biohorizon (cf. callomon 1984), i.e. as period is the time equivalent of system. biohorizons are the smallest palaeontologically correlatable segments of geological time using ‘guide fossils’ and unlike ‘normal’ chronozones are effectively defined at both their bases and tops. their duration is typically geologically very short but a significant time gap may exist between each successive unit and is shown as an interval on any correlation diagram (page 1992, 1995a; dommergues et al. 1994a). the use of biohorizons is rather like events in event stratigraphy, as they enable the correlation of virtually isochronous time lines between successions at different localities (callomon 1984, 1985). by convention, zonules are quoted in a similar fashion to zones and subzones with a non-italicised specific name (e.g. planorbis zonule) but biohorizons typically retain an italicised specific epithet (e.g. planorbis biohorizon). biohorizonal and zonule schemes have been derived in different regions for different parts of the jurassic and relevant lower jurassic schemes are introduced below. they represent the ultimate in resolvable chronology for the jurassic, and the average zonule or biohorizon plus interval duration is potentially less than 200 000 years in the lower jurassic of northwest europe (page 1995a). ammonoid provincialism in the lower jurassic of europe: consequences for correlation ammonoids, like any group of organisms, frequently show distinctive geographical distribution patterns, reflecting ecological and physical controls on individuals and populations. such patterns are characterised as biogeographical provinces and the inevitable consequence of using ammonites for correlation purposes is that every province, almost by definition, will have a different scheme of standard zones. these differences can inevitably make interprovincial correlations at zonal, and especially subzonal and horizon level, difficult. up to four 26 contemporaneous faunal provinces can be recognised in the lower jurassic of europe. most are geographically adjacent and inter-provincial correlations are thus usually fairly good, although the southernmost faunas of the mediterranean province are sometimes sufficiently distinct as to present significant correlation problems when compared to better known successions further north. as discussed by page (1996), the main faunal provinces recognised in the lower jurassic of europe are: 1. northwest european province. the province was characteristic of much of europe from the hettangian to the early pliensbachian and again in the late toarcian, when great faunal uniformity characterised most of the region. faunal affinities are virtually entirely from southern or tethyan areas (i.e. the mediterranean province) and direct connection to the arctic or boreal sea was lacking (except perhaps in the earlier late toarcian). zonal schemes of the northwest european province are well-developed reflecting the long history of research on the area. at times in the pliensbachian, and also earlier, faunas in the lusitanian basin (portugal) developed a separate character from those of the rest of the northwest european province (dommergues & mouterde 1987), thereby creating some correlation problems. a separate sequence of ‘horizons’ has consequently been recognised for this area for the lower pliensbachian, but whether full province or simply sub-province status is warranted is unclear. 2. mediterranean province. this was characteristic of deeper water areas of southern and south-eastern europe (italy, austria, southern spain, etc.) throughout the jurassic. faunal sequences of the province are often less well-known in detail than those of more northerly areas, and zonations therefore tend to be relatively crude in comparison, but with considerable potential for refinement. 3. subboreal province. the establishment of direct marine connections with the boreal sea in the late pliensbachian and toarcian enabled some mixing of arctic province faunas and the previously separated faunas of northwest europe (see fig. 2). the abundance of boreal taxa alongside more southerly forms distinguishes the province in the more northerly areas of europe (such as northern britain). the zonation is well-established and correlates fairly well with more southerly areas as a result of faunal overlap. 4. submediterranean province. the province developed in parallel with the subboreal province in the late pliensbachian in regions between the former and the mediterranean province (e.g. in mid to southern france). boreal taxa are usually infrequent and an independent zonation is used in the lower toarcian, but with significant similarities to that of subboreal areas. the following sections summarise and correlate available zonal schemes for each of the faunal provinces or sub-provinces recognised in europe. the emphasis is on the northwest european and subboreal schemes as these are most appropriate for northern european areas (including britain, northern france, northern germany and denmark). stratotypes are only identified when clearly described in the literature. it would be premature to propose such definitions for other units, pending a reassessment of taxal ranges and surviving localities. only the taxa which are useful for correlating each subzone are cited; other forms will be present but these are not always chronologically diagnostic. hettangian stage the hettangian stage as originally proposed by renevier (1864) corresponded to the first two zones of the schemes proposed by oppel (1856–1858) for the jurassic, namely those of ammonites [psiloceras] planorbis and ammonites [schlotheimia] angulatus. this interpretation remains essentially unchanged, the only difference being the later creation of a liasicus ‘zone’ for the lower part of the original angulatus zone. the base of the hettangian stage and the jurassic system the jurassic colloquium in luxembourg in 1962 recommended that the planorbis ‘zone’ should form the lowest part of the hettangian stage (mauberge 1964); this zone, as conventionally interpreted, marks the first occurrence of ammonites in northwest europe, after the re-establishment of fully marine conditions towards the end of the triassic period. the type locality of the index fossil of the first subchronozone and the first chronozone of the jurassic of the northwest european province (i.e. psiloceras planorbis (j. de c. sowerby 1824)) is on the coast of west somerset near watchet in south-west england. a proposal 27 in 1967 to the 2nd luxembourg colloquium on the jurassic system recommended that a type section for the zone should actually be selected in this region (d.t. donovan, p.e. kent and h.c. ivimey-cook in: morton 1971). there has been much subsequent discussion as to where exactly the boundary should be drawn (e.g. torrens & getty 1980; warrington & ivimey-cook 1990), culminating in the proposal of warrington et al. (1994) to establish a section at st audries bay, east of watchet as a gssp. the issue as to whether the region is in fact suitable for such a definition has not, however, been thoroughly addressed (page 1994; page et al. 1994). the latter proposal placed the base of the subchronozone and chronozone at the then first recorded occurrence of ammonites in bed a21 of palmer (1972; equivalent to beds 13–15 of whittaker & green 1984). ammonites had not previously been recorded any lower in the immediate area and, indeed, characteristically triassic ammonoids are entirely lacking in britain (although there is a problematic record of an indeterminate, and therefore completely undiagnostic psiloceratid, from presumed latest triassic deposits elsewhere in the region; donovan et al. 1989). this definition, therefore, relies to a certain extent on negative evidence, as emphasised by the subsequent discovery of earlier ammonites at the same locality by hodges (1994) in beds a18 and a19 (beds 8 and 9 of whittaker & green 1984) and the author, westwards along the coast in doniford bay nearer watchet. further work on the sections has revealed a succession of ammonite faunas, previously unrecognised, which correlates well with faunas recovered from the wilkesley borehole in cheshire in north-west england (bloos & page 1997, 2000a; page & bloos 1998). the lowest fauna in the borehole is characterised by the ribbed psiloceratid, psiloceras erugatum (phillips), as already noted by d.t. donovan (in: poole & whiteman 1966, pp. 50, 140). re-examination of the higher faunas reveals the presence of neophyllites and the subchronozonal index p. planorbis itself. the same sequence is present in somerset with bed 8 now known to yield p. erugatum (bloos & page 1997, 2000a; page & bloos 1998). no ammonites are presently known from demonstrably lower levels in britain. in response to the new discoveries in somerset, warrington & ivimey-cook (1995) modified their original proposal and placed the base of the jurassic system at the base of bed a18 (bed 8 of whittaker & green 1984). as discussed by bloos & page (1997, 2000a), however, elsewhere in europe the erugatum fauna has not yet been positively identified, the earliest ammonites typically belonging to neophyllites (bloos 1999) or p. ex grp sampsoni (portlock)/psilonotum (quenstedt), indicating a degree of diachroneity in the first occurrence of ammonites. even where late triassic choristoceras ammonite faunas are present below psiloceratids, such as in mediterranean regions, there remains an ‘ammonite gap’, and the earliest ‘jurassic’ ammonites also appear to represent later species close to the p. sampsoni/psilonotum group (bloos 1985; hallam 1990). more complete sequences of ammonoid faunas from the uppermost triassic (rhaetian stage) to the lowermost jurassic are known, however, elsewhere in the world and two have been proposed as candidate gssps in new york canyon, nevada, usa (guex 1980, 1982; guex et al. 1997) and northern peru (von hillebrandt 1994, 1997). whether a ‘new world’ definition for the base of the jurassic system is acceptable on historical grounds remains to be determined, but it is clear that at the present state of knowledge, it is not possible to accurately correlate these sections with any in europe (bloos & page 2000a). the st audries bay section remains the best exposed triassic–jurassic boundary section in britain, although the ammonite faunas are better preserved in doniford bay along the somerset coast to the west. the erugatum fauna itself, however, is very rare and poorly preserved in somerset, but much better developed, both in terms of abundance and preservation, in cheshire and north yorkshire in northern england, although only known in situ in boreholes (bloos & page 2000a). none of these localities is, therefore, ideal as a gssp, not least due to the absence of triassic ammonoids below in a continuous open marine sequence, but also due to the present lack of clear records of p. erugatum outside britain. if the latter could be identified elsewhere, however, for instance amongst some of the early and poorly characterised psiloceratids in new york canyon, a final agreement on the selection of a suitable gssp for the base of the jurassic system would be much closer. hettangian correlative schemes and ammonoid provincialism hettangian faunas are remarkably similar globally, reflecting the relatively small number of available ammonoid taxa, so soon after the late triassic mass extinctions. no direct boreal links existed in europe, so a simple pattern of a northern northwest european province and a southern and deeper water mediterranean province is recognisable (fig. 1). 28 29 northwest european province (britain, ireland, france, germany, etc.) the origins of the standard zonation for the hettangian of the region go back to the original scheme presented by oppel (1856–1858) with the later addition of a liassicus zone by collenot (1869). the zonation as presently employed is that established by d.t. donovan (in: dean et al. 1961) with later additions by elmi & mouterde (1965) and bloos (1979, 1983), summarised diagrammatically by mouterde & corna (1991) and reviewed by mouterde & corna (1997). mouterde & corna (1991, 1997) presented a scheme of zonules which are here integrated with a sequence of biohorizons established as a result of recent studies in south-west england by page (1994, 1995b, 2002a), page & bloos (1998) and bloos & page (2000a, b). few of the units summarised below have established or proposed stratotypes, with the notable exception of the planorbis subchronozone (and therefore the planorbis chronozone) at the base of the stage and hence the jurassic system. cox (1990) has suggested reference sections for each zone, although these are not equivalent to basal boundary stratotypes; some of these proposals would require revision, however, based on new information on the succession of ammonite faunas both at the proposed localities and elsewhere. planorbis chronozone index. psiloceras planorbis (j. de c. sowerby 1824). author. oppel (1856). planorbis subchronozone index. as planorbis chronozone (see above). author. trueman (1922). stratotype. proposed by warrington et al. (1994), modified by warrington & ivimey-cook (1995); base of bed 8 (= a18), st audries bay, west somerset, england (see discussion above). correlating fauna. dominated by smooth whorled species of ‘psiloceras’ and neophyllites (sensu lange pseudomoreana depressa striatissima complanata similis extranodosa amblygonia hadroptychus laqueolus laqueus schroederi portlocki hagenowi prometheus intermedium johnstoni caloceras sp.2 caloceras sp.1 plicatulum sampsoni α planorbis antecedens imitans erugatum marmorea/depressa ? extranodosa ? storthoceras/ alsatites, etc. tenerum calliphyllum ‘psilonotum’ marmorea megastoma calliphyllum depressa complanata extranodosa hadroptychus liassicus laqueus portlocki belcheri johnstoni plicatulum sampsoni planorbis depressa complanata extranodosa laqueus portlocki johnstoni planorbis angulata liassicus planorbis chronozone subchronozone zonule biohorizon faunas chronozone northwest european province mediterranean province fig. 1. hettangian subdivisions and correlations: northwest european and mediterranean provinces. for explanation, see text. the double lines separating the biohorizons on this and later figures indicate the stratigraphic interval conceptually present between each successive division of this kind (see discussion in text). 1941) but with ribbed psiloceras erugatum at the base of the subchronozone. included subdivisions. planorbis zonule (including psiloceras erugatum, neophyllites imitans, n. antecedens and psiloceras planorbis biohorizons), sampsoni [= psilonotum] zonule (including psiloceras sampsoni α biohorizon), plicatulum zonule (including plicatulum biohorizon) (mouterde & corna 1991, 1997; page 1994; page & bloos 1998; bloos & page 2000a, b). comment. buckman (1930 in: buckman 1909–1930) first proposed a plicatus hemera, subsequent raised to the status of ‘horizon’ (= zonule) by elmi & mouterde (1965) and ultimately a subzone by von hillebrandt (1990). status as a zonule is provisionally retained here, however, pending an assessment of ammonite faunas in south-west england, where circumstantial evidence suggests that at least three, maybe four, correlatable biohorizons could ultimately be recognisable, potentially lending support to raising the status of the unit to subchronozone. johnstoni subchronozone index. caloceras johnstoni (j. de c. sowerby 1824). author. von schloenbach (1863) as a zone, trueman (1922) as a subzone (but see discussion by d.t. donovan in: dean et al. 1961, p. 444). correlating fauna. early species of caloceras. included subdivisions. johnstoni zonule (including caloceras sp. 1, caloceras sp. 2 and c. johnstoni biohorizons), belcheri zonule (including c. intermedium biohorizon) (mouterde & corna 1991, 1997; page 1994). liassicus chronozone index. alsatites liassicus (d’orbigny 1844). author. collenot (1869). portlocki subchronozone index. waehneroceras portlocki (wright 1881). author. lang (1924). correlating fauna. early species of waehneroceras including w. (curviceras), also psilophyllites. included subdivisions. portlocki zonule (including w. (curviceras) prometheus, psilophyllites hagenowi, waehneroceras portlocki and w. schroederi biohorizons) (mouterde & corna 1991, 1997; page 2002a). laqueus subchronozone index. laqueoceras laqueus (quenstedt 1856). author. reynès (1879). correlating fauna. laqueoceras is typical, with alsatites above, in association with common late waehneroceras spp. (= ‘saxoceras’) and late caloceras at higher levels. included subdivisions. laqueus zonule (including laqueoceras laqueus biohorizon), liassicus zonule (including alsatites laqueolus biohorizon), hadroptychus zonule (including w. hadroptychus biohorizon) (mouterde & corna 1991, 1997). angulata chronozone index. schlotheimia angulata (schlotheim 1820). author. oppel (1856). extranodosa subchronozone index. schlotheimia extranodosa (waehner 1886). author. lange (1922) as germanica zone (name changed due to apparent synonymy by donovan 1952). correlating fauna. early species of schlotheimia. included subdivisions. extranodosa zonule (including schlotheimia amblygonia and s. extranodosa biohorizons) (bloos 1984; mouterde & corna 1991, 1997). complanata subchronozone index. schlotheimia complanata von koenen 1902. author. spath (1942) replacing the stenorhyncha zone of lange (1922). correlating fauna. schlotheimia spp. including relatively compressed and smooth whorled macroconchs, with giant forms at higher levels. the rare early arietitid schreinbachites is locally recorded. included subdivisions. complanata zonule (including s. similis, s. complanata and s. striatissima biohorizons) (bloos 1984; mouterde & corna 1991, 1997; page 1995b; bloos & page 2000b). depressa subchronozone index. schlotheimia depressa (waehner 1886). author. bloos (1983). correlating fauna. the index species and its allies, especially s. ex grp pseudomoreana (spath). ?paracaloceras is also known, but is very rare. included subdivisions. depressa zonule (including s. depressa and s. pseudomoreana biohorizons) (bloos & page 2000b; page 2002a). comment. the depressa subchronozone was established by bloos (1983, 1984) but was reduced in status to a zonule by mouterde & corna (1991, 1997). in britain, the subchronozone is well-developed, although has only recently been recognised; it 30 includes the pseudomoreana and ‘schlotheimia sp. 2’ biohorizons of page (1995b) (page & bloos 1998; bloos & page 2000b). mediterranean province (austria) the subdivision of the hettangian mediterranean province is not well-developed despite the presence of considerably richer ammonoid faunas than in northwest european areas. a sequence of three zones, calliphyllum, megastoma and marmorea, was recognised by waehner (1886) and used again by lange (1952). blind (1963) subdivided the calliphyllum ‘zone’ and attempted to subdivide the higher zones, although bloos (1984) considered that these schemes were not entirely satisfactory. the marmorea chronozone has been the subject of some discussion as to its meaning and indeed whether it ranges into the sinemurian (taylor 1986), but in a european sense at least, a wholly hettangian age is most likely (bloos 1983, 1984). calliphyllum chronozone index. psiloceras calliphyllum neumayr 1879. author. waehner (1886). correlating fauna. species of psiloceras, especially the index, and early waehneroceras (= curviceras) at higher levels. included subdivisions. the psilonotum and calliphyllum faunas (= zones i–ii of blind 1963; bloos 1984). megastoma chronozone index. ?kammerkarites megastoma (waehner 1886). author. waehner (1886). correlating fauna. species of kammerkarites, waehneroceras (sensu lato), alsatites and caloceras. included subdivisions. the tenerum, storthoceras/alsatites and extranodosa faunas (blind 1963; bloos 1984). marmorea chronozone index. schlotheimia marmorea (oppel 1856). author. waehner (1886). correlating fauna. species of schlotheimia with common early arietitidae (paracaloceras, etc.). included subdivisions. includes an s. marmorea/s. depressa fauna at the top of the chronozone (as indicated by bloos 1984). sinemurian stage the sinemurian stage corresponds to the zonal range suggested by oppel (1856) minus the hettangian of renevier (1864), i.e. with a bucklandi chronozone at the base and a raricostatum chronozone at the top. additional zones have been added subsequently, but the modern interpretation remains essentially unchanged. although the lower sinemurian substage has no common alternative name, the term lotharingian (from ‘lorraine, france’) is often used in france for the upper sinemurian (after haug 1910 in: haug 1908–1911 but sensu spath 1942 as haug’s stage included the turneri zone, now considered to be lower sinemurian; dean et al. 1961). the base of the sinemurian stage the base of the stage is drawn at the base of the conybeari subchronozone, the lowest division of the bucklandi chronozone. d.t. donovan (in: morton 1971) proposed that the stratotype for the stage should be established on the dorset–devon coast, near lyme regis. in this area, the relatively well-preserved and first obvious sinemurian fauna (i.e. dominated by arietitid ammonites, including metophioceras) occurs in nodules on the base of bed 19 of lang (1924) although traces of vermiceras have now been found about 0.25 m lower, in bed 18 (author’s unpublished data), and 0.25 m above the last schlotheimia seen on the surface of bed 17 (page 1992); this is c. 0.75 m lower than donovan’s original designation of the base of bed 21 for the base of the stage. recent study of other sections on the coast of the neighbouring county of somerset has revealed a considerably expanded hettangian–sinemurian succession, more than four times thicker than that at lyme regis (palmer 1972; whittaker & green 1984; page 1992, 1995b). in particular, it has been possible to demonstrate that near lyme regis the faunal record is incomplete (page 1992, 1995b; bloos & page 2000b). the dorset section is not therefore a suitable stratotype let alone a candidate gssp, whereas the somerset section has great potential and was proposed as such by page et al. (2000) and ratified by iugs at the rio de janeiro symposium in august 2000 – the first lower jurassic gssp to be formally approved (bloos & page 2002). it is the faunal completeness and expanded succession (c. 14 m for the earliest sinemurian conybeari subzone alone) which makes the locality unusual in europe. elsewhere, for 31 instance in germany and south-east france, successions are usually much thinner and much less complete. the base of the sinemurian at east quantoxhead is taken at the first occurrence of abundant arietitid ammonites 70 cm above the base of bed c100/145 (palmer 1972; whittaker & green 1984) and only 20 cm above the last common hettangian schlotheimia. this fauna, with vermiceras quantoxense (bloos & page) and v. palmeri (bloos & page), underlies an assemblage including metophioceras ex grp brevidorsale (quenstedt) and m. conybearoides (reynes) which is normally the earliest recorded in the sinemurian elsewhere in europe (bloos 1997; bloos & page 2000a; page et al. 2000). sinemurian correlative schemes and ammonoid provincialism sinemurian provincialism is essentially the same as in the hettangian, with a broad northwest european province over much of europe, and a mediterranean province in south-easternmost areas. some endemism or geographic restriction of genera and species in portugal at times in the late sinemurian probably indicates some links with north africa (i.e. the ethiopian province sensu page 1996; figs 2, 3). northwest european province (britain, ireland, france, germany, switzerland, etc.) the subdivision of the stage into chronozones and subchronozones, as started by oppel (1856–1858) and developed through the work of w.d. lang and l.f. spath (lang et al. 1923; lang 1924; spath 1924, 1942; lang & spath 1925b), was eventually stabilised by d.t. donovan (in: dean et al. 1961). the only subsequent changes have been in the name of zonal indices as a result of the identification of senior synonyms. ivimeycook & donovan (1984) proposed the removal of the bucklandi subchronozone and lowering of the base of the semicostatum zone to include most of the former subzone. page (1992), however, retained a bucklandi subchronozone as the unit forms a significant and important stratigraphical unit that is readily recognisable internationally. corna et al. (1997) provided a recent review of the zonation of the stage and the zonules therein, essentially reproducing the scheme presented diagrammatically by corna et al. (1991). bloos (1985) proposed a sequence of ‘horizons’ for the basal part of the stage, the divisions identified being very similar to those of the earliest part of a zonule scheme described by corna (1987). page (1992) introduced a sequence of biohorizons for the entire sinemurian, the upper sinemurian part of which was correlated with french successions in burgundy (dommergues 1993) to produce the revised biohorizonal scheme of dommergues et al. (1994a). additional biohorizons in the lower sinemurian, especially the conybeari subchronozone were recognised by page (1995a, b), bloos & page (2000b) and page et al. (2000). reference sections for each zone were suggested by cox (1990), but now require some modification based on new faunal information from various regions. lower sinemurian substage bucklandi chronozone index. arietites bucklandi (j. sowerby 1818). author. oppel (1858). conybeari subchronozone index. metophioceras conybeari (j. sowerby 1816). author. tutcher (1918) as a zone, trueman (1922) as a subzone. stratotype. level 0.7 m above the base of bed c100/145, east quantoxhead, somerset, uk (see discussion above). correlating fauna. species of metophioceras, with early vermiceras and epammonites. giant charmasseiceras typical. included subdivisions. latisulcatum zonule (including vermiceras quantoxense, metophioceras sp. 2, and m. conybearoides biohorizons), rotarium zonule (including epammonites rotarius, m. rouvillei, coroniceras rotator and v. elegans biohorizons), conybeari zonule (including metophioceras conybeari biohorizon) (corna 1987; corna et al. 1991; page 1992, 1995a, b; bloos & page 2000b, 2002; page et al. 2000). rotiforme subchronozone index. coroniceras rotiforme (j. de c. sowerby 1824). author. collenot (1879) as a zone, trueman (1922) as a subzone. correlating fauna. evolute and strongly ribbed species of coroniceras, with occasional charmasseiceras at certain levels. included subdivisions. hyatti zonule (including epammonites silvestrei, coroniceras cf. defneri, c. rotiforme, c. aff. rotiforme and c. caprotinum biohorizons), 32 33 ‘schloenbachi’ zonule (including c. aff. kridion biohorizon) (corna 1987; corna et al. 1991, 1997; page 1992, 1995b). bucklandi subchronozone index. as bucklandi chronozone (see above). author. spath (1942) as a subzone. correlating fauna. large and massive whorled arietitidae (arietites) typical with coroniceras at certain levels. early arnioceras and charmasseiceras present, the former in the upper part of the subchronozone. 31: cf. bordoti 30: subturneri 29: birchi 28: pseudobonnardi 27: obtusiformis 26: hartmanni 25: brooki 24: sulcifer 23: cf. semicostatum 22: alcinoeiforme 21: euagassiceras 20: cf. resupinatim 19: pseudokridion 18: acuticarinatum 17b: alcinoe 17a: paracoroniceras sp 16: bodleyi 15b: cf. charlesi 15a: lyra 14: multicostatum 13: cf. scunthorpense 12: isis 11: aff. isis 10: scylla 9: kridion 8: caprotinum 7b: aff. rotiforme 7a: rotiforme 6: cf. defneri 5c: silvestrei 5b: conybeari 5a: elegans 4: rotator 3b: rouvillei 3a: rotarius 2b: conybearoides 2a: metophioceras sp.b 1: quantoxense bordoti turneri brooki sauzeanum nodulatum scipionanum alcinoe crossi charlesi lyra bisulcatus isis coronaries schloenbachi hyatti conybeari rotarium ‘latisulcatum’ birchi brooki sauzeanum scipionanum lyra bucklandi rotiforme conybeari turneri semicostatum bucklandi chronozone subchronozone zonule biohorizon ceratoides ? part rotiforme fauna northwest european province mediterranean province ? fig. 2. lower sinemurian subdivisions and northwest european horizons. for explanation, see text. included subdivisions. coronaries zonule (including vermiceras scylla biohorizon, previously included in the rotiforme subzone by page 1992 but placed at the base of the bucklandi subchronozone by corna et al. 1997), isis zonule (including arietites aff. isis, a. isis and a. aff. scunthorpense biohorizons), bisulcatus zonule (including coroniceras multicostatum biohorizon) (corna 1985; corna et al. 1991, 1997; page 1992, 1995b). semicostatum chronozone index. arnioceras semicostatum (young & bird 1829). author. judd (1875). lyra subchronozone index. paracoroniceras lyra (hyatt 1867). author. buckman (1918) as gmuendense zone, changed to reynesi subzone by d.t. donovan (in: dean et al. 1961, after a junior synonym of p. lyra, subchronozone renamed accordingly by guérin-franiatte 1966). equivalent to the charlesi subzone of mouterde & tintant (1980). correlating fauna. species of paracoroniceras and ‘pararnioceras’, with common arnioceras at certain levels. included subdivisions. lyra zonule (including coroniceras lyra (sensu stricto) biohorizon), charlesi zonule (including p. cf. charlesi [= c. lyra in page 1992] and arnioceras bodleyi biohorizons), crossi zonule (probably includes the paracoroniceras sp. biohorizon of page 1992), alcinoe zonule (including the ‘pararnioceras’ alcinoe biohorizon; included in the scipionanum subzone by corna 1987 and corna et al. 1991, 1997) (corna 1987; corna et al. 1991; page 1992). scipionanum subchronozone index. agassiceras scipionanum (d’orbigny 1844). author. tutcher (1918). correlating fauna. species of agassiceras and arnioceras. may include early euagassiceras (e. striaries (quenstedt)) in the higher part of the subchronozone (teste corna et al. 1997). included subdivisions. scipionanum zonule (including arnioceras acuticarinatum biohorizon), nodulatum zonule (including arnioceras pseudokridion biohorizon) (corna 1987; corna et al. 1991, 1997; page 1992). resupinatum subchronozone index. euagassiceras resupinatum (simpson 1843). author. tutcher (1918) as a zone, trueman (1922) as a subzone. correlating fauna. euagassiceras and arnioceras. included subdivisions. sauzeanum zonule (including e. cf. resupinatum, ‘euagassiceras’ sp., ‘pararnioceras’ alcinoeiforme and arnioceras cf. semicostatum biohorizons) (corna et al. 1991, 1997; page 1992). turneri chronozone index. caenisites turneri (j. de c. sowerby 1824). author. wright (1860). brooki subchronozone index. caenisites brooki (j. sowerby 1818). author. lang (1914, w.d. lang in: lang et al. 1923). correlating fauna. early species of caenisites and some arnioceras and ‘sulciferites’ (i.e. late charmasseiceras). included subdivisions. brooki zonule (including sulciferites sulcifer, c. brooki and arnioceras hartmanni biohorizons) (corna et al. 1991, 1997; page 1992). birchi subchronozone index. microderoceras birchi (j. sowerby 1820). author. collenot (1869) as a zone, spath (1942) as a subzone. correlating fauna. species of microderoceras, caenisites ex grp turneri, early epophioceras locally and common early promicroceras at higher levels. included subdivisions. turneri zonule (including caenisites obtusiformis, epophioceras pseudobonnardi, microderoceras birchi and caenisites subturneri biohorizons), bordoti zonule (including caenisites cf. bordoti biohorizon) (corna et al. 1991, 1997; page 1992). upper sinemurian substage obtusum chronozone index. asteroceras obtusum (j. sowerby 1817). author. oppel (1856). obtusum subchronozone index. as obtusum chronozone (see above). author. spath (1942) as a subzone of ‘promicroceras planicosta or obtusum (sensu stricto)’ (= planicosta zone of lang 1914). correlating fauna. early species of asteroceras, with promicroceras (but not p. planicosta), xipheroceras and occasionally late arnioceras. 34 35 tardecrescens meigeni boehmi crassicostatum raricostatum rhodanicum edmundi delicatum oxynotum gagateum aff. glaber fowleri blakei stellare confusum apianatum macdonelli raricostatum densinodulum oxynotum simpsoni denotatus stellare obtusum raricostatum oxynotum obtusum xl: aplanatum/tardecrescens xxxix: recticostatum xxxviii: aureolum xxxvii: macdonnelli xxxvi: meigeni xxxv: subplicatum xxxiv: boehmi xxxili: cf. intermedium xxxil: crassicostatum xxxi: raricostatum xxx: rhodanicum xxix: ‘echioceras’ sp.3 xxviii: radiatum xxvii: grp armatum xxvi: bispinigerum xxv: lymense xxiv: subplanicosta xxiii: delicatum xxii: doris xxi: bifer xx: grp oxynotum xix: driani xviii: gagateum xvii: exortum xvi: aff. glaber xv: denotatus xiv: fowleri xiii: cf. undaries xii: sagittarium xi: aff. arnouldi x: blakei s.s. ix: stellare viii: cf. landrioti vii: margaritoides vi: aff. margaritoides v: ‘galaticeras’ iv: obtusum iii: semicostatoides ii: cf. confusum i: aff. confusum tardecrescens/romanicum oosteri/gruenae meigeni/macdonnelli meigeni meigeni/charpentieri liciense favrei boehmi raricostatoides quenstedti rigidum salisburgense glaber eparietites suevicum ceratoides (part) chronozone subchronozone zonule biohorizon biohorizon northwest european province mediterranean province ? grp oxynotum aff. saltriense retusum fig. 3. upper sinemurian (‘lotharingian’) subdivisions and northwest european horizons. for explanation, see text. included subdivisions. confusum zonule (including asteroceras aff. confusum, as. cf. confusum, arnioceras semicostatoides and as. obtusum biohorizons) (corna et al. 1991, 1997; page 1992; dommergues et al. 1994a). as. confusum is here preferred to as. obtusum as an index for a zonule, as the latter is relatively rarely recorded and the former is more typical of this level in northwest europe. stellare subchronozone index. asteroceras stellare (j. sowerby 1815). author. buckman (1910) as a zone, lang (1914) as a subzone. correlating fauna. late species of asteroceras, also promicroceras, xipheroceras, some epophioceras with early aegasteroceras at the top of the subchronozone. included subdivisions. stellare zonule (including ‘galaticeras’ [=? bouhamidoceras], asteroceras aff. margaritoides, as. margaritoides, epophioceras cf. landrioti and as. stellare biohorizons), blakei zonule (including aegasteroceras blakei, arnioceras aff. arnouldi and aegasteroceras sagittarium biohorizons) (corna et al. 1991, 1997; page 1992; dommergues et al. 1994a). denotatus subchronozone index. eparietites denotatus (simpson 1855). author. buckman (1918) as a hemera; buckman (1919 in: buckman 1909–1930) in a zonal context. correlating fauna. species of eparietites, with aegasteroceras (grp simile) at lower levels. included subdivisions. fowleri zonule (including eparietites cf. undaries, ep. fowleri and ep. denotatus biohorizons), glaber zonule (including ep. aff. glaber biohorizon) (corna et al. 1991, 1997; page 1992; dommergues et al. 1994a). oxynotum chronozone index. oxynoticeras oxynotum (quenstedt 1843). author. oppel (1856). simpsoni subchronozone index. oxynoticeras simpsoni (simpson 1843). author. buckman (1918) as a hemera, spath (1942) as a subzone. correlating fauna. o. ex grp simpsoni and gagaticeras. included subdivisions. gagateum zonule (including gagaticeras exortum, g. gagateum and paroxynoticeras driani biohorizons) (corna et al. 1991, 1997; page 1992; dommergues et al. 1994a). oxynotum subchronozone index. as oxynotum chronozone (see above). author. buckman (1918) as a hemera, spath (1942) as an alternative name for his bifer subzone. correlating fauna. oxynoticeras grp oxynotum, bifericeras and some palaeoechioceras. included subdivisions. oxynotum zonule (including o. ex grp oxynotum, bifericeras ex grp bifer and gleviceras doris biohorizons) (corna et al. 1991, 1997; page 1992; dommergues et al. 1994a). raricostatum chronozone index. echioceras raricostatum (zieten 1831). author. oppel (1856). densinodulum subchronozone index. crucilobiceras densinodulum buckman 1923. author. buckman (1923 in: buckman 1909–1930) as a hemera. w.d. lang (1926 in: lang & spath 1926) as a subzone. correlating fauna. crucilobiceras, eoderoceras, gleviceras and at restricted levels, early echioceratids (including ‘plesechioceras’ [= ‘paltechioceras’ sensu lato] etc.). included subdivisions. delicatum zonule (including ‘plesechioceras’ delicatum biohorizon), edmundi zonule (including crucilobiceras subplanicosta, oxynoticeras lymense, eoderoceras bispinigerum and eoderoceras ex grp armatum biohorizons), radiatum zonule nov. (including echioceras radiatum and echioceras [= paltechioceras sensu lato] sp. 3 biohorizons) (corna et al. 1991, 1997; page 1992; dommergues et al. 1994a). raricostatum subchronozone index. as raricostatum chronozone (see above). author. buckman (1918) as a raricostatoides hemera, spath (1942) as a subzone (e. raricostatoides is generally considered to be a junior synonym of e. raricostatum). correlating fauna. raricostate echioceras spp., with some crucilobiceras, eoderoceras and gleviceras. included subdivisions. rhodanicum zonule (including echioceras rhodanicum biohorizon), raricostatum zonule (including e. raricostatum biohorizon), crassicostatum zonule (including e. crassicostatum biohorizon), boehmi zonule (including ‘paltechioceras’ cf. intermedium and ‘p.’ boehmi biohorizons) (corna et al. 1991, 1997; page 1992; dommergues et al. 1994a). 36 macdonnelli subchronozone index. leptechioceras macdonnelli (portlock 1943). author. buckman (1918) as a hemera, w.d. lang (1926 in: lang & spath 1926) as a zone, spath (1942) as a subzone. correlating fauna. leptechioceras spp., with some eoderoceratids (including epideroceras in more southerly areas) and early radstockiceras. included subdivisions. meigeni zonule (including leptechioceras subplicatum, l. meigeni and l. macdonnelli biohorizons) (corna et al. 1991, 1997; page 1992; dommergues et al. 1994a). aplanatum subchronozone index. paltechioceras aplanatum (hyatt 1889). author. buckman (1918) as a hemera, lang (1926) as a zone, spath (1942) as a subzone. correlating fauna. late species of paltechioceras, with eoderoceras and epideroceras, the latter typically in more southerly areas. included subdivisions. tardecrescens zonule (including paltechioceras aureolum, p. rectiradiatum and p. aplanatum/tardecrescens biohorizons) (corna et al. 1991, 1997; page 1992; dommergues et al. 1994a). comment. the eoderoceras donovani and vicininodiceras simplicicosta biohorizons were provisionally included in the sinemurian by page (1992) and dommergues et al. (1994a). new information from north yorkshire, however, suggests that at least the former is best considered to be of basal pliensbachian age. the latter remains stratigraphically problematic, however, and is consequently no longer used as a biohorizonal index, pending clarification of the precise position of the specimens recorded by donovan (1990) on raasay, scotland. mediterranean province (austria, italy) the basic northwest european chronozonal and subchronozonal framework is generally used in mediterranean areas, for example by donovan (1990), dommergues et al. (1995), blau (1998) and especially blau & meister (1999). at zonule/biohorizonal level, however, correlation is typically less precise as many species appear to be different and the relative generic composition is also often distinct in mediterranean areas. in the lower sinemurian, only generalised correlations seem to be possible (corna et al. 1991, 1997), but in the upper sinemurian, there are considerably more links, and in the raricostatum chronozone in particular, some biohorizons are even correlatable between provinces, as demonstrated by blau & meister (1999; fig. 3). summarising from figures 2–4 of blau & meister (1999), the following is a generalised and provisional composite sequence of biohorizons or faunas for the mediterranean province in the upper sinemurian: 1. obtusum chronozone, obtusum subchronozone (confusum zonule, may include elements of the arnioceras ceratoides fauna of the apennines, italy); stellare subchronozone (stellare zonule, including asteroceras retusum and as. aff. saltriense (part?) faunas/biohorizons (adnet, austria); blakei zonule, including as. aff. saltriense (?part) (adnet) and as. suevicum (lienz, austria) faunas/biohorizons); denotatus subchronozone (fowleri zonule, including eparietites fauna/biohorizon (lienz, austria); glaber zonule (including e. glaber fauna/biohorizon (lienz and adnet, austria)). 2. oxynotum chronozone, simpsoni subchronozone (gagateum zonule, no recorded faunas in blau & meister 1999), oxynotum subchronozone (oxynotum zonule, including oxynoticeras ex grp oxynotum (adnet, austria) and paroxynoticeras salisburgense (lienz, austria) faunas/biohorizons). 3. raricostatum chronozone, densinodulum subchronozone (delicatum–edmundi zonules, part?, including g. rigidum fauna/biohorizon, part?); raricostatum subchronozone (rhodanicum zonule, including echioceras quenstedti biohorizon; raricostatum zonule, including e. raricostatoides biohorizon; crassicostatum zonule, no confirmed records in blau & meister 1999; boehmi zonule, including paltechioceras boehmi biohorizon (apennines, italy)); macdonelli subchronozone (meigeni zonule, including ‘paltechioceras’ favrei, ‘p.’ liciense, leptechioceras meigeni/p. charpentieri, l. meigeni and l. meigeni/l. macdonnelli biohorizons (lienz, austria)); aplanatum subchronozone (tardecrescens zonule, including paltechioceras oosteri/miltoceras gruenae (lienz, austria) and p. tardecrescens/p. romanicum biohorizons (apennines, italy and lienz, austria)). localities and faunas are described by dommergues et al. (1994b; apennines, italy), dommergues et al. (1995; adnet, austria) and blau (1998; lienz, austria). 37 pliensbachian stage the stage name was first used by oppel (1856) as ‘pliensbach-gruppe’, with a zonal composition identical to that still used throughout europe. pliensbach is near boll in württemberg, southern germany. oppel’s stage is essentially the same as d’orbigny’s earlier nongeographically named (and hence rejected) ‘liasien’ (d’orbigny 1842). ‘charmouthian’ is another early synonym attributed to mayer-eymar (1864) but considered by dean et al. (1961) as having been first published by renevier (1874). unlike other lower jurassic stages, the pliensbachian is often subdivided into named substages, the lower pliensbachian corresponding to the carixian substage (after ‘carixa’ = charmouth, dorset, england; lang 1913) and the upper pliensbachian corresponding to the domerian substage (after monte domaro in the lombardy alps, italy; bonarelli 1894). the base of the pliensbachian stage the first zone of oppel’s original ‘pliensbach-gruppe’ is still used as the first of the stage, namely the jamesoni ‘zone’ (oppel 1856), now with a taylori subchronozone forming its lowest unit. d.t. donovan (in: morton 1971) discussed the base of the stage and indicated that the taylori subzone was first recognised on the coast of southern england in dorset near charmouth. the base of the stage at that locality (= base of bed 105 of lang 1928) lies immediately above a non-sequence which omits the highest two subzones of the sinemurian. the locality is not suitable, therefore, for defining the base of the stage, according to ics guidelines. a non-sequence at the base of the stage is also widespread throughout much of northern europe, including at pliensbach itself (geyer 1964; morton 1971). where more complete sinemurian–pliensbachian successions are exposed, however, for instance in robin hood’s bay (north yorkshire, england; tate & blake 1876; dommergues & meister 1992; page 1992; hesselbo & jenkins 1995) and the isle of raasay (western scotland; oates 1978; donovan 1990; page 1992), drawing a suitable base can be somewhat problematic. in robin hood’s bay, above the last typical sinemurian-type ammonites (paltechioceras) is a fauna with a small eoderoceratid described by dommergues & meister (1992) as ‘bifericeras’ donovani. at this level occur the first, rare apoderoceras, a genus more characteristic of the taylori subzone than the index phricodoceras taylori itself, but it is not until slightly higher in the yorkshire succession that the latter genus is recorded. on raasay, however, above the last paltechioceras and below common apoderoceras, donovan (1990) reported a fauna with a rare liparoceratid, vicininodiceras. this fauna was included in the sinemurian by donovan and also by page (1992), but its assignment to the stage is somewhat problematic, especially as its stratigraphical relationship to ‘bifericeras’ donovani is presently unknown. recent work on the robin hood’s bay sections by the pliensbachian working group of the isjs (in 1999) has confirmed the presence of small (?juvenile) apoderoceras in association with ‘bifericeras’ donovani (= donovani biohorizon or zonule), that this level is therefore suitable for marking the base of the pliensbachian stage and that the fauna is typical of a conventional pliensbachian and not the terminal sinemurian as implied by dommergues et al. (1994a, 1997). the section has now yielded results from other stratigraphical techniques, including microfossil assemblages and chemostratigraphy, leading to a formal proposal as a gssp in late 2002 (meister et al. 2002), the base of the stage being drawn at the base of bed 73 (or 1011) of dommergues & meister (1992). pliensbachian correlative schemes and ammonoid provincialism early pliensbachian faunas show a great deal of uniformity throughout northern europe and most of the region is included in a northwest european province. mediterranean province faunas, although differing in taxonomic detail, are sufficiently well-correlated with those further north and west that the same standard zonation is used for both provinces, but faunal successions are nevertheless distinguishable at horizonal level (figs 4, 5). in the upper part of the lower pliensbachian and throughout the upper pliensbachian, however, the establishment of direct connections with boreal regions resulted in a faunal spectrum developing across europe from assemblages dominated by boreal taxa in the northern areas (characterising a subboreal province) through faunas dominated by mediterranean faunas in central and western areas (submediterranean province) to true mediterranean province faunas in the south. no true arctic province faunas (sensu page 1996) are recognisable in europe, however. inter-provincial faunal links facilitate good correlations and it is possible to use the 38 same standard zonation throughout europe, but faunal sequences may be very different at horizon level. the development of endemism in the early pliensbachian of portugal has necessitated the establishment of a different sequence of horizons from adjacent submediterranean areas and sub-provincial or full provincial status may therefore be justifiable. northwest european province (britain, france, germany, northern spain, etc.) the zonal and subzonal framework of the province is used throughout europe and is based on oppel’s original scheme (oppel 1856). for convenience, the province is here equated temporally with the lower pliensbachian substage (equivalent to the first three zones of oppel’s scheme) and a description of the upper pliensbachian zonation of northern areas is included under the heading of subboreal province. the subzonal framework follows d.t. donovan (in: dean et al. 1961) as reviewed by dommergues et al. (1997). reference sections for each chronozone were proposed by cox (1990) but some need modification in the light of new information on faunal successions elsewhere. a number of systems of horizons, primarily in the sense of zonules, have been proposed for the early part of the stage in the province and include the schemes of dommergues (1979) and phelps (1985), with later revisions by dommergues & meister (1987), dommergues (1987, 1997) and dommergues et al. (1991). lower pliensbachian substage jamesoni chronozone index. uptonia jamesoni (j. de c. sowerby 1827). author. oppel (1856). taylori subchronozone index. phricodoceras taylori (j. de c. sowerby 1826). author. spath (1923). stratotype. proposed by meister et al. (2002): base of bed 73/1011, robin hood’s bay, north yorkshire, uk (see discussion above). correlating fauna. species of apoderoceras, with rarer phricodoceras and radstockiceras. tetraspidoceras is common in southern areas (dommergues et al. 1997). included subdivisions. donovani zonule, nodogigas/grp aculeatum zonule, taylori zonule (dommergues et al. 1991, 1997; meister et al. 2002). comment. the basal unit of the stage characterised by bifericeras donovani is here quoted as a ‘zonule’ for consistency with higher subdivisions of the pliensbachian. in meister et al. (2002), however, the unit is treated as a ‘biohorizon’. polymorphus subchronozone index. polymorphites polymorphus (quenstedt 1845). author. buckman (1918) as a hemera, spath (1923) as a subzone. correlating fauna. species of polymorphites, with early platypleuroceras in the upper part of the subchronozone. radstockiceras and epideroceras also locally present. included subdivisions. polymorphus zonule (dommergues et al. 1991, 1997). brevispina subchronozone index. platypleuroceras brevispina (j. de c. sowerby 1827). author. von seebach (1864) as part of a larger zone, buckman (1918) as a hemera, spath (1923) as a subzone. correlating fauna. platypleuroceras ex grp brevispina and spp. radstockiceras and tragophylloceras also present, with metaderoceras in more southerly areas. included subdivisions. brevispina zonule, submuticum zonule (dommergues et al. 1991, 1997). jamesoni subchronozone index. as jamesoni chronozone (see above). author. buckman (1918) as a hemera, d.t. donovan (in: dean et al. 1961) as a subzone. correlating fauna. uptonia ex grp jamesoni and spp. tragophylloceras and coeloceras also typical locally. included subdivisions. jamesoni zonule, pettos zonule (dommergues et al. 1991, 1997). ibex chronozone index. tragophylloceras ibex (quenstedt 1843). author. oppel (1856). masseanum subchronozone index. tropidoceras masseanum (d’orbigny 1844). author. buckman (1918) as a hemera, spath (1923) as a subzone. correlating fauna. tropidoceras spp., tragophylloceras frequent, especially in more southerly areas. included subdivisions. masseanum zonule (dommergues 1979; phelps 1985; dommergues et al. 1991, 1997). 39 valdani subchronozone index. acanthopleuroceras valdani (d’orbigny 1844). author. collenot (1869) as a zone, buckman (1918) as a hemera, spath (1942) as a subzone. correlating fauna. acanthopleuroceras spp., with tragophylloceras, liparoceras, and beaniceras at higher levels. included subdivisions. arietiforme zonule, maugenesti zonule, valdani zonule, actaeon zonule, alisiense zonule (dommergues 1979; phelps 1985; dommergues et al. 1991, 1997). luridum subchronozone index. beaniceras luridum (simpson 1855). 40 figulinum angulatum crescens capricornus lataecosta maculatum sparsicosta luridum crassum rotundum alisiense actaeon valdani maugenesti arietiforme masseanum pettos jamesoni submuticum brevispina polymorphus taylori nodogigas/grp aculeatum donovani figulinum capricornus maculatum luridum valdani masseanum jamesoni brevispina polymorphus taylori dilectum ? demonense aenigmaticum davoei ibex jamesoni beirense amaltheiforme splendens polymorphoides renzi maugenesti carinatum uptonia sp. bronni–lata bronni–jamesoni acanthobronnii muellensis costatus biruga dayiforme caprariforme chronozone subchronozone northwest european province mediterranean province zonule (nw european s.s.) zonule (lusitanian) zonule chronozone (spain) d ay ic er oi de s costicillatum volubile–pantanellii dilectum aff. dilectum catriense ? mediterraneum flandrini erythreum ? sellae ? fig. 4. lower pliensbachian (‘carixian’) subdivisions and correlations: northwest european and mediterranean provinces. for explanation, see text. author. d.t. donovan (in: dean et al. 1961). correlating fauna. beaniceras grp luridum with liparoceras spp. and tragophylloceras. included subdivisions. rotundum zonule, crassum zonule, luridum zonule (dommergues 1979; phelps 1985; dommergues et al. 1991, 1997). davoei chronozone index. prodactylioceras davoei (j. sowerby 1822). author. oppel (1856). maculatum subchronozone index. androgynoceras maculatum (young & bird 1822). author. lang (1936) as a zone, spath (1938) as a subzone. correlating fauna. a. ex grp maculatum and liparoceras. included subdivisions. sparsicosta zonule, maculatum zonule (dommergues 1979; phelps 1985; dommergues et al. 1991, 1997). capricornus subchronozone index. androgynoceras capricornus (schlotheim 1820). author. wright (1863) used the index for a zone equivalent to the davoei zone of oppel; retained as a subzone by d.t. donovan (in: dean et al. 1961). correlating fauna. a. lataecosta, a. grp capricornus, etc. with liparoceras (including l. (becheiceras)) and, at certain levels, prodactylioceras. included subdivisions. lataecosta zonule, capricornus zonule, crescens zonule (dommergues 1979; phelps 1985; dommergues et al. 1991, 1997). figulinum subchronozone index. oistoceras figulinum (simpson 1855). author. lang (1936). correlating fauna. oistoceras spp. with liparoceras (becheiceras), tragophylloceras and prodactylioceras at certain levels. included subdivisions. angulatum zonule, figulinum zonule (dommergues 1979; phelps 1985; dommergues et al. 1991, 1997). northwest european province: lusitanian ?sub-province (portugal) faunal differentiation in the early pliensbachian has resulted in some correlation difficulties at intra-subzonal level between portuguese faunas and those elsewhere in europe. general affinities are however northwest european. local subdivisions as tabulated by dommergues et al. (1991, 1997, based on dommergues 1987) are shown in figure 4, and include: 1. jamesoni chronozone, taylori subchronozone (unnamed interval, caprariforme zonule, dayiforme zonule); polymorphus subchronozone (biruga zonule, costatus zonule); brevispina subchronozone (muellensis zonule, acanthobronni zonule); jamesoni subchronozone (bronni–jamesoni zonule, bronni– lata zonule, uptonia sp. zonule – ?part). 2. ibex chronozone, masseanum subchronozone (uptonia sp. zonule – ?part); valdani subchronozone (dayiceroides zonule, including acanthopleuroceras carinatum and a. maugenesti ?biohorizons, renzi zonule, polymorphoides zonule, splendens zonule, amaltheiforme zonule, beirense zonule and unnamed interval); luirdum subchronozone and davoei chronozone as northwest european province. subboreal province (britain, northern germany, etc.) increased faunal polarisation between southern and northern pliensbachian faunas in the upper pliensbachian (= domerian substage) makes intra-subzonal correlations difficult or impossible. subboreal faunas in particular, are dominated by amaltheidae with only very rare representation of mediterranean hildocerataceae. the former dominate in the south, however, and correlate detailed horizonal sequences. the basic zonal/subzonal framework employed is nevertheless identical to that used elsewhere in europe and follows m.k. howarth (in: dean et al. 1961). significantly, howarth (1991–1992) has proposed definitions for the bases of the subzones of the substage using northern english sites, all falling within the scope of the subboreal province as interpreted here. no complete system of biohorizons has been specifically proposed for the province, but the sequence of zonules used by dommergues et al. (1997) can be applied. full descriptions of faunas and successions can be found in howarth (1955, 1956, 1957, 1991–1992). upper pliensbachian substage margaritatus chronozone index. amaltheus margaritatus (de montfort 1808). author. oppel (1856). 41 stokesi subchronozone index. amaltheus stokesi (j. sowerby 1818). author. lang (1936). defined base. base of bed 1, hawsker bottoms or bed 12, staithes, staithes sandstone formation, north yorkshire, england (howarth 1955, 1991–1992; cox 1990). correlating fauna. amaltheus ex grp stokesi with rare protogrammoceras. tragophylloceras and liparoceras (becheiceras) present, especially in more southerly areas. included subdivisions. occidentale zonule, monestieri zonule, nitescens zonule, celebratum zonule (dommergues 1979; phelps 1985; dommergues et al. 1991, 1997). comment: the zonules of dommergues et al. (1997) are recognised largely on the basis of protogrammoceras which is typically very rare or absent in more northerly areas, thereby making the recognition of such units increasingly difficult from south to north. subnodosus subchronozone index. amaltheus subnodosus (young & bird 1828). author. spath (1942) as a nodifer subzone; name changed to subnodosus by howarth (1955) for reasons of synonym. defined base. base of bed 18, hawsker bottoms or bed 26, staithes, cleveland ironstone formation, north yorkshire, england (howarth 1955, 1991–1992). correlating fauna. amaltheus ex grp subnodosus with rare protogrammoceras. included subdivisions. normanianum zonule, unnamed interval (dommergues et al. 1991, 1997). comment: the normanianum zonule of dommergues et al. (1997) is recognised on the basis of protogrammoceras which is typically very rare or absent in more northerly areas, thereby making recognition of the unit increasingly difficult from south to north. 42 hawskerense lotti solare transiens salebrosum ? ruthense algovianum bertrandi (il) kurrianus ugdulenai macrum ragazzonii boscense depressum celebratum nitescens monestieri occidentale elisa emaciatum imitator solare levidorsatum meneghinii ? accuratum canavarii ugdulenai bertrandi (l) ragazzonii cornacaldense celebratum marianii isseli/ brevispiratum lavinianum/ portisi elisa solare levidorsatum meneghenii accuratum bertrandi ragazzonii cornacaldense emaciatum algovianum lavinianum hawskerense elaboratum solare transiens salebrosum (unnamed interval) algovianum (unnamed interval) normanianum celebratum nitescens monestieri occidentale hawskerense apyrenum gibbosus subnodosus stokesi spinatum margaritatus chronozone subchronozone zonule zonule zonule (austria–italy–spain) subchronozone (spain) chronozone (spain) mediterranean province submediterranean province subboreal province isseli brevispiratum lavinianum portisi fig. 5. upper pliensbachian (‘domerian’) subdivisions and correlations: subboreal, submediterranean and mediterranean provinces. for explanation, see text. gibbosus subchronozone index. amaltheus gibbosus (schlotheim 1820). author. buckman (1918) as a hemera, kuhn (1935) as a zone, spath (1942) as a subzone. defined base. base of bed 21, hawsker bottoms or bed 32, staithes, cleveland ironstone formation, north yorkshire, england (howarth 1955, 1991–1992). correlating fauna. amaltheus ex grp gibbosus, and rare protogrammoceras. included horizons. unnamed interval, algovianum zonule, unnamed interval. comment. the algovianum zonule of dommergues et al. (1997) is recognised on the basis of protogrammoceras which is typically very rare or absent in more northerly areas, thereby making the recognition of the unit increasingly difficult from south to north. spinatum chronozone index. pleuroceras spinatum (bruguière 1789). author. oppel (1856). apyrenum subchronozone index. pleuroceras apyrenum (buckman 1911). author. spath (1942). defined base. base of bed 25, hawsker bottoms, cleveland ironstone formation, north yorkshire (howarth 1955, 1991–1992; non cox 1990). correlating fauna. pleuroceras ex grp solare (phillips), p. apyrenum, etc. and occasional amaltheus ex grp margaritatus. included subdivisions. salebrosum zonule, transiens zonule, solare zonule (dommergues et al. 1991, 1997). hawskerense subchronozone index. pleuroceras hawskerense (young & bird 1928). author. buckman (1922) as a hemera, spath (1942) as a subzone. defined base. base of bed 38, hawsker bottoms, or bed 32, staithes, cleveland ironstone formation, north yorkshire (howarth 1955, 1991–1992). correlating fauna. pleuroceras ex grp hawskerense and occasional pseudamaltheus engelhardti (d’orbigny). included subdivisions. elaboratum zonule, hawskerense zonule (dommergues et al. 1991, 1997). submediterranean province (southern france, northern spain, etc.) the submediterranean province developed in parallel with the subboreal and is also therefore primarily a phenomenon of the upper pliensbachian (= domerian). faunas are dominated by hildocerataceae, with some dactylioceratidae but amaltheidae are usually much less common than in subboreal areas. in consequence, a separate sequence of zonules is recognised for the upper pliensbachian of the submediterranean province (meister 1989; dommergues et al. 1991, 1997) as follows: 1. margaritatus chronozone, stokesi subchronozone (as subboreal province); subnodosus subchronozone (depressum zonule, boscense zonule); gibbosus subchronozone (ragazzonii zonule, macrum zonule, ugdulenai zonule, kurrianus zonule, bertrandi zonule (ii), algovianum zonule, ruthense zonule, unnamed interval). 2. spinatum chronozone, apyrenum subchronozone (as subboreal province); hawskerense subchronozone (lotti zonule, hawskerense zonule). mediterranean province (austria, italy, southern spain) faunas in the province are commonly so different from those of northern europe that it is remarkable that the same standard zonation can often be used. the presence of interprovincial correlative tie-lines is very important in linking northern and southern areas, but many problems still remain, leading to the introduction by braga et al. (1982) and braga (1983) of a new zonal scheme for the mediterranean province, as correlated by dommergues et al. (1997) with submediterranean and northwest european schemes and reproduced here on figures 4 and 5. several sequences of ‘horizons’ are also recognised within the mediterranean province for the pliensbachian stage and schemes for austria/italy and spain are also tabulated on figures 4 and 5 (after braga 1983; meister 1987; ferreti 1990; dommergues et al. 1991, 1997; meister et al. 1994). toarcian stage the stage name was derived by d’orbigny (1842–1849) from the town of thouars in central western france (poitou) where the division is well-developed and rich in ammonites. unlike those of the earlier lower jurassic stages, all the zones of the modern toarcian stage postdate oppel’s simple scheme of 1856 (with a ‘zone der 43 posidonia bronni’, followed by a ‘zone des ammonites jurensis’). as for the pliensbachian, the toarcian is divided into two substages. both these substages have names, the lower toarcian corresponding to the whitbian (after whitby in north yorkshire, north-east england; buckman 1910) and the upper toarcian to the yeovilian (after yeovil in somerset, south-west england; buckman 1910). the original definition of whitbian included the variabilis zone but was redefined by m.k. howarth (in: dean et al. 1961) to exclude that zone which was reassigned as the lowest of the yeovilian. unlike for the pliensbachian, however, these geographical names for substages have fallen into disuse. the base of the toarcian stage the base of the lower zone of oppel’s toarcian, (with the bivalve ‘posidonia’ bronni as index) corresponds to the changeover in northwest europe from typical pliensbachian ammonite faunas with pleuroceras to typical toarcian faunas with abundant dactylioceras. the base of the stage in the type region (submediterranean province as used here) corresponds to the base of the tenuicostatum chronozone. this zone, however, has its type locality on the coast of northern england (north yorkshire; buckman 1910, subboreal province), and this has led to various proposals or assumptions that this latter area should include the basal boundary stratotype of the stage (e.g. m.k. howarth in: morton 1971; cox 1990; howarth 1991–1992). the lower part of the tenuicostatum chronozone as used in both provinces is a paltum subchronozone, the base of which was defined by howarth (1991–1992) as the base of bed 26, kettleness, or the base of bed 58, staithes (basal grey shales member, whitby mudstone formation), both in north yorkshire (sections as described by howarth 1955, 1973). the use of this stratigraphic division dates back to buckman (1922) with reference to a fauna in the highly condensed middle–upper lias ‘junction bed’ of dorset, southern england (buckman 1910; jackson 1926). the characteristic protogrammoceras fauna of buckman’s paltus hemera is abundant in dorset, but does not occur as part of a clear faunal succession due to the highly condensed and lenticular nature of the uppermost pliensbachian – lower toarcian ‘junction bed’. in yorkshire, however, the fauna is much rarer, although the rock succession itself is considerably expanded. nevertheless, as noted by howarth (1992, p. 5 in: howarth 1991–1992), p. paltum is probably not a good index, but was selected as it is restricted to this level in britain. further south in europe, similar protogrammoceras (or ‘paltarpites’) occur at higher and lower levels (howarth 1992, p. 7 in: howarth 1991–1992). in addition, the earliest toarcian dactylioceras, in association with p. paltum, are abundant and characteristic in southern areas but virtually absent in britain. the association of the former, including dactylioceras (eodactylites) mirabilis fucini and d. (e.) simplex fucini, with p. paltum in spain (e.g. goy & martinez 1990) and even in southern germany (schlatter 1985) presents a better scenario, however, for defining the base of the stage (cf. elmi 1997; elmi et al. 1997). detailed correlations between the northern european tenuicostatum chronozone and the unit of the same name in southern areas is not yet established in detail, due to apparent differences in the dactylioceratid faunas (cf. elmi et al. 1997). nevertheless, it is very unlikely that sections in britain will be suitable candidate gssps as they are faunally impoverished when compared to southern europe and hence have a more limited international correlation potential. toarcian correlative schemes and ammonoid provincialism early toarcian faunas show similar distribution patterns to those of the late pliensbachian, as boreal connections appear to have persisted throughout the substage. as in the pliensbachian, a subboreal to submediterranean to mediterranean province transition is recognisable. unlike the pliensbachian, however, separate zonal schemes exist for all three provinces (fig. 6). there are very close links, however, between submediterranean and subboreal faunas, so whether a distinct zonal/subzonal scheme is needed is debatable. even at intra-subzonal level, similarities are sufficient that many cross-correlations are possible. in the late toarcian, these similarities are so great that only one zonal scheme is justifiable in northwest europe and only one scheme is therefore reviewed here for a northwest european province (fig. 7). rare boreal links include occasional pseudolioceras, mainly in northern britain. this unified scheme combines elements of the british standard of w.t. dean (in: dean et al. 1961) and howarth (1980) with the french schemes of gabilly et al. (1971, 1974), gabilly (1976) and elmi et al. (1991, 1994, 1997). 44 subboreal province (northern britain) a restricted subboreal province is most characteristic of northern england and scotland in the lower toarcian (from late tenuicostatum to bifrons chronozones) when faunas are dominated by dactylioceratids with less common hildocerataceae, especially in scotland. the presence at certain levels of boreal hildocerataceans, such as tiltoniceras, elegantuliceras, ovaticeras and pseudolioceras, is typical. further south in britain, faunas become more submediterranean in character and late dactylioceratids of the bifrons chronozone are rare in dorset and somerset. the following scheme for the lower toarcian is that of m.k. howarth (in: dean et al. (1961) as modified by howarth (1973) and defined by howarth (1991–1992). the included biohorizons are described more fully elsewhere (page 2002b) and are based on observations by the author and records by howarth (1962, 1973, 1978, 1991–1992). the system of zonules for the ‘north west european’ province of elmi et al. (1997; essentially the submediterranean province as used here) is applicable, at least in part, to the subboreal province and is utilised below. lower toarcian substage tenuicostatum chronozone index. dactylioceras tenuicostatum (young & bird 1822). author. buckman (1910) replacing the annulatus zone of tate & blake (1876) (the latter with a stratigraphically incorrectly placed index; see m.k. howarth in: dean et al. 1961). paltum subchronozone index. protogrammoceras paltum (buckman 1922). author. buckman (1922) as a hemera, howarth (1973) as a subzone. defined base. provisionally defined (see comment) as base of bed 55, staithes, ketteness member, cleveland ironstone formation, north yorkshire, england (howarth 1955, 1973, 1991–1992; cox 1990). correlating fauna. protogrammoceras paltum, and very rare early dactylioceras (including d. pseudocommune fucini). included subdivisions. paltum zonule (including protogrammoceras paltum biohorizon). comment. although useful for correlating the base of the toarcian in a subboreal/uk sense, howarth’s north yorkshire reference section is not a ratified gssp (see discussion above). 45 crassum–semipolitum crassum–bifrons vortex braunianus turriculatum athleticum commune ovatum falciferum pseudoserpentinum elegans exaratum elegantulum antiquum semicelatum tenuicostatum clevelandicum crosbeyi paltum crassum fibulatum commune falciferum exaratum semicelatum (l) tenuicostatum clevelandicum paltum bifrons serpentinum tenuicostatum bifrons serpentinum tenuicostatum bifrons levisoni polymorphum bifrons sublevisoni falciferum? levisoni ‘semicelatum’ (il) mirabile semipolitum bifrons apertum lusitanicum tethysi sublevisoni striatus paltarpites bifrons sublevisoni falciferum elegantulum ‘semicelatum’ (ll) paltum semipolitum bifrons apertum lusitanicum tethysi sublevisoni douvillei pseudoserpentinum strangewaysi elegantulum ‘semicelatum’ tenuicostatum crosbeyi paltum chronozone subchronozone biohorizon zonule subchronozone chronozone chronozone subchronozone zonule subboreal province submediterranean province mediterranean province fig. 6. lower toarcian subdivisions and correlations: subboreal, submediterranean and mediterranean provinces. for explanation, see text. clevelandicum subchronozone index. dactylioceras clevelandicum howarth 1973. author. howarth (1973) = dactylioceras sp. nov. horizon of m.k. howarth (in: dean et al. 1961). defined base. base of bed 18, grey shales member, whitby mudstone formation, north yorkshire coast (e.g. kettleness/port mulgrave), england (howarth 1973, 1991–1992). correlating fauna. d. grp clevelandicum. included subdivision. crosbeyi zonule (including dactylioceras crosbeyi and d. clevelandicum biohorizons). tenuicostatum subchronozone index. as tenuicostatum chronozone (see above). author. buckman (1930 in: buckman 1909–1930) as a hemera, m.k. howarth (in: dean et al. 1961) as a ‘horizon’, howarth (1973) as a subzone. defined base. base of bed 20, grey shales member, whitby mudstone formation, north yorkshire coast (kettleness/port mulgrave), england (howarth 1973, 1991–1992). correlating fauna. d. tenuicostatum. included subdivisions. tenuicostatum zonule (including dactylioceras tenuicostatum biohorizon). semicelatum (i) subchronozone index. dactylioceras semicelatum (simpson 1843). author. m.k. howarth (in: dean et al. 1961) as ‘d. semicelatum horizon’, howarth (1973) as a subzone. the index has been used by mauberge (1948, 1952) for a tenuicostatum–semicelatum zone and a semicelatum zone, effectively as a synonym of the tenuicostatum chronozone as a whole. defined base. base of bed 28, grey shales member, whitby mudstone formation, north yorkshire coast (e.g. kettleness/port mulgrave area), england (howarth 1973, 1991–1992). correlating fauna. d. semicelatum and tiltoniceras antiquum (wright). included subdivisions. semicelatum zonule (including dactylioceras semicelatum and tiltoniceras antiquum biohorizons). comments. the semicelatum subchronozone of the subboreal province is as restricted by howarth (1973). in the submediterranean and mediterranean provinces, however, a ‘semicelatum subzone’ is also used but broadly equates to the clevelandicum, tenuicostatum and semicelatum subchronozones combined. clearly, different names are needed to distinguish the two units, but pending a full revision the two subchronozones are here distinguished as semicelatum (i) and semicelatum (ii), respectively. serpentinum chronozone index. harpoceras serpentinum (schlotheim 1813). author. oppel (1856) suggested the use of ammonites serpentinum as an alternative index to posidonia bronni. reynès (1868) restricted a serpentinus zone to the lower part of the lower toarcian (below a bifrons zone) and its use was restricted further by the recognition of a basal tenuicostatum zone. the use of h. serpentinum as an index has priority over h. falciferum (j. sowerby 1820) first used as a zonal index by haug (1885) and subsequently by most british authors (e.g. m.k. howarth in: dean et al. 1961; howarth 1991–1992). exaratum subchronozone index. cleviceras exaratum (young & bird 1828). author. buckman (1910) as a zone, arkell (1956) as a subzone. defined base. base of bed 33, mulgrave shale member, whitby mudstone formation, north yorkshire coast (e.g. rosedale wyke to lingrow knock or saltwick bay area, whitby), england (howarth 1962, 1991–1992; cox 1990). correlating fauna. elegantuliceras, cleviceras and early harpoceras in succession. dactylioceras (nodicoeloceras) and hildaites ex grp levisoni present. included subdivisions. elegantulum zonule (including elegantuliceras elegantulum biohorizon), strangewaysi zonule (including cleviceras exaratum and c. elegans biohorizons). falciferum subchronozone index. harpoceras falciferum (j. sowerby 1820). author. haug (1885) as a broader zone (including the exaratum subzone), buckman (1930 in: buckman 1909–1930) as a hemera, arkell (1933) as a subzone. defined base. base of bed 41, mulgrave shale member, whitby mudstone formation, north yorkshire coast (e.g. saltwick bay area, rosedale wyke to lingrow knock, etc.) england (howarth 1962, 1991–1992). correlating fauna. harpoceras ex grp falciferum. included subdivisions. pseudoserpentinum zonule (including harpoceras pseudoserpentinum biohorizon), douvillei zonule (including h. falciferum biohorizon). bifrons chronozone index. hildoceras bifrons (bruguière 1789). author. reynès (1868) as a zonal index. 46 commune subchronozone index. dactylioceras commune (j. sowerby 1815). author. wright (1863) as a zone in a broad sense, and equivalent to lower toarcian, m.k. howarth (in: dean et al. 1961) as a restricted subzone. equivalent to the ‘communis beds’ or subcarinata zone of thompson (1910). suggested re-defined base. base of bed 47, whitby, alum shale member, whitby mudstone formation, north yorkshire, england (sections described by howarth 1962, 1991–1992). this definition is modified from that of howarth (1991–1992) and cox (1990) to include the ovatum biohorizon. it is suggested that this modification is necessary to equate the base of the bifrons chronozone in the subboreal province with that in the submediterranean and mediterranean provinces where faunas with hildoceras ex grp sublevisoni are taken to indicate the lower part of the chronozone (abundant h. ex grp sublevisoni occurs with rare ovaticeras ovatum near ilminster in southern england; unpublished observations by the author). correlating fauna. dactylioceras, including d. commune and d. athleticum (simpson), also ovaticeras. hildoceras spp., including h. lusitanicum meister are usually uncommon. included subdivisions. sublevisoni zonule (including ovaticeras ovatum biohorizon), tethysi zonule (including dactylioceras commune biohorizon), lusitanicum zonule (including d. athleticum biohorizon). fibulatum subchronozone index. peronoceras fibulatum (j. de c. sowerby 1823). author. thompson (1910). defined base. base of bed 60, alum shale member, whitby mudstone formation, whitby, north yorkshire, england. correlating fauna. peronoceras, zugodactylites and porpoceras spp. with some hildoceras (including h. bifrons). included subdivisions. apertum zonule (including peronoceras turriculatum and zugodactylites braunianus biohorizons), bifrons zonule (part, including porpoceras vortex biohorizon). crassum subchronozone index. catacoeloceras crassum (young & bird 1828). author. corroy & gérard (1933). defined base. 1.5 m above the base of bed 72, whitby or the base of bed xiiv, alum shale member, whitby mudstone formation, ravenscar, north yorkshire, england (howarth 1962, 1978, 1991–1992). correlating fauna. abundant catacoeloceras with less frequent hildoceras including h. semipolitum buckman. included subdivisions. bifrons zonule (part, including catacoeloceras crassum – h. bifrons biohorizon), semipolitum zonule (including c. crassum – h. semipolitum biohorizon). submediterranean province (southern england, france, germany, northern spain, etc.) although strong links exist between northern and southern regions, faunas of submediterranean areas in the lower toarcian are usually richer in hildocerataceae, sometimes to the virtual exclusion of dactylioceratidae. the separate chronozonal schemes historically used in britain and france are not entirely necessary as the greatest faunal changes actually occur within britain, between southern england (dorset and somerset) and northern england (north yorkshire) and scotland (isle of raasay). the chronozonal framework employed in france is identical in zonal composition to that in britain, but differs at the level of subzones and horizons due to faunal differentiation. the scheme summarised here follows elmi et al. (1991, 1994, 1997) and is largely based on gabilly et al. (1971) and gabilly (1976). lower toarcian substage tenuicostatum chronozone index and author. as subboreal province (see above). paltum (or paltus) subchronozone index, author and defined base. as subboreal province (see above). correlating fauna. protogrammoceras paltum and, in certain areas, dactylioceras (eodactylites) spp., neolioceratoides and lioceratoides. included subdivisions. paltum [paltus] zonule (gabilly et al. 1974; gabilly 1976; elmi et al. 1991, 1994, 1997). semicelatum (ii) subchronozone index and author. as subboreal province, but used here in a less restricted sense (sensu gabilly et al. 1971; see comments below). 47 correlating fauna. dactylioceras spp. including d. semicelatum, and rarer hildoceratids, including protogrammoceras and neolioceratoides. included subdivisions. crosbeyi zonule, tenuicostatum zonule, semicelatum zonule (gabilly 1976; elmi et al. 1991, 1994, 1997). comments. the exact relationship of the submediterranean semicelatum subchronozone to the subboreal subchronozone of the same name is problematic as both the crosbeyi and tenuicostatum subchronozones are difficult to recognise outside their type area. the conventional usage continued here is to draw the base of the submediterranean division at the first occurrence of abundant dactylioceras (orthodactylites) – this would therefore equate with the base of the clevelandicum subchronozone of subboreal areas. as the definition of the subchronozone in that region has been stabilised by formal definition (howarth 1991–1992, see above), it is necessary to rename the submediterranean division. as suggested above in discussion of the subchronozone in the subboreal province, pending full revision the two units are here distinguished as ‘semicelatum’ (i) (subboreal) and ‘semicelatum’ (ii) (submediterranean and mediterranean). serpentinum chronozone index and author. as subboreal province (see above). elegantulum subchronozone index. elegantuliceras elegantulum (young & bird 1828) author. gabilly (1976) as a subzone, equivalent to his earlier strangewaysi zone. correlating fauna. elengantuliceras, harpoceras ex grp serpentinum, cleviceras, hildaites and dactylioceratids including ‘nodicoeloceras’. included subdivisions. elegantulum zonule, strangewaysi zonule (gabilly 1976; elmi et al. 1991, 1994, 1997). comment. the subchronozone is probably exactly equivalent to the subboreal exaratum subchronozone as it appears to have a coincident base (the definition of howarth 1991–1992 cited above could also therefore apply to this subchronozone; indeed, this is necessary if the serpentinus chronozone is to be used in both provinces). falciferum subchronozone index and author. as subboreal province (see above). correlating fauna. harpoceras ex grp falciferum, hildaites and dactylioceratids including nodicoeloceras. included subdivisions. pseudoserpentinum zonule, douvillei zonule (gabilly 1976; elmi et al. 1991, 1994, 1997). comments. the subchronozone as used in submediterranean areas is probably equivalent to that of subboreal regions. a common defined base is therefore necessary if the same name is to be used in both provinces and that of howarth (1991–1992) may be appropriate. bifrons chronozone index and author. as subboreal province (see above). sublevisoni subchronozone index. hildoceras sublevisoni fucini 1922. author. donovan (1958). correlating fauna. abundant hildoceras spp. including hi. ex grp sublevisoni, harpoceras ex grp falciferum, dactylioceras ex grp commune, etc. included subdivisions. sublevisoni zonule, tethysi zonule, lusitanicum zonule (gabilly 1976; elmi et al. 1991, 1994, 1997). comments. see discussion on subboreal province commune subchronozone regarding equilibration with the submediterranean scheme. howarth (1992, p. 177 in: howarth 1991–1992) includes faunas with hi. sublevisoni within a falciferum subchronozone sensu anglico. as harpoceras ex grp falciferum persists into the sublevisoni subchronozone, however, this interpretation would conflict with the established interpretation of french authors in the type region of the bifrons chronozone (e.g. gabilly et al. 1971) where the first appearance of primitive hildoceras (i.e. hi. ex grp sublevisoni) is regarded as marking the base of the zone. this latter convention is followed here. bifrons subchronozone index and author. as bifrons chronozone (see above). correlating fauna. hildoceras bifrons and allied species abundant, with rarer harpoceras subplanatum (oppel), phymatoceras and dactylioceratids, including zugodactylites and catacoeloceras. included subdivisions. apertum zonule, bifrons zonule, semipolitum zonule (gabilly 1976; elmi et al. 1991, 1994, 1997). comments. a semipolitum subchronozone has been used in submediterranean province areas (e.g. in elmi et al. 1991, 1994), after donovan (1958), but was reduced to the status of ‘horizon’ (= zonule) by elmi et al. (1997). 48 49 northwest european province (britain, france, germany, northern spain, etc.) in the upper toarcian, good faunal links throughout most of more northerly areas of europe suggest that only one zonal scheme for the region is justifiable (fig. 7). faunas are richer in more southerly areas, especially in phymatoceratidae, but the bulk of the correlatively important grammoceratinae (hildoceratidae) are very widespread. the following scheme is based on the french systems of gabilly et al. (1971, 1974) and elmi et al. (1991, 1994, 1997), taking into account the earlier british zonation of w.t. dean (in: dean et al. 1961) as modified by howarth (1980). all intrasubzonal units are here used as zonules, no complete schemes of biohorizons presently being available for the province. upper toarcian substage variabilis chronozone index. haugia variabilis (d’orbigny 1845). author. buckman (1888 in: buckman 1887–1907) as a subzone of larger jurensis zone. first used as a zone by welsch (1903). variabilis subchronozone index. as variabilis chronozone (see above). author. buckman (1925 in: buckman 1909–1930) as a hemera, used as a horizon by gabilly et al. (1971) and a subzone by gabilly (1976). correlating fauna. haugia ex grp variabilis, denckmannia and catacoeloceras dumortieri (de brun). included subdivisions. navis zonule, jugosa zonule (gabilly 1976; elmi et al. 1991, 1994, 1997). comment. interpretation of the base of the variabilis chronozone differs between britain and france depending on whether the last occurrence of hildoceras semipolitum/catacoeloceras spp. or the first occurrence of haugia is taken as a guide (the former would indicate bifrons chronozone, crassum subchronozone in britain and the latter variabilis chronozone, variabilis subchronozone in france). the base of the conventional variabilis chronozone is therefore drawn at a lower level in france than in britain as all three taxa are recorded together. this latter convention is followed here. illustris subchronozone index. haugia illustris (denckman 1887). buckmani lugdunensis celtica mactra tectiforme pseudoradiosa munieri insignisimilis gruneri pachu cappucinum fallaciosum fascigerum thouarsense doerntense bingmanni vitiosa phillipsi illustris jugosa navis aalensis pseudoradiosa dispansum thouarsense variabilis fluitans mactra pseudoradiosa levesquei gruneri insigne fallaciosum (l) fascigerum striatulum bingmanni vitiosa illustris variabilis ? fluitans flamandi mactra meneghinii subregale aratum aalensis meneghinii speciosum bonarellìi gradata chronozone chronozonesubchronozonesubchronozone mediterranean provincenorthwest european province zonule zonule fluitans mactra ? reynesi speciosum fallaciosum (ll) mediterraneum alticarinatus ? gemma fig. 7. upper toarcian subdivisions and correlations: northwest european and mediterranean provinces. for explanation, see text. author. stolley (1909) as a zone, gabilly et al. (1971) as a horizon, gabilly (1976) as a subzone. correlating fauna. haugia grp illustris, h. phillipsi (simpson), etc. (including haugiella), hammatoceras spp., denckmannia. included subdivisions. illustris zonule, phillipsi zonule (gabilly 1976; elmi et al. 1991, 1994, 1997). vitiosa subchronozone index. haugia vitiosa (buckman 1909). author. gabilly (1976) as a subzone. correlating fauna. haugia grp vitiosa, denckmannia, hammatoceras, pseudogrammoceras. included subdivisions. vitiosa zonule (gabilly 1976; elmi et al. 1991, 1994, 1997). thouarsense chronozone index. grammoceras thouarsense (d’orbigny 1844). author. brasil (1896) as a zone. comment. in britain, only two subchronozones are conventionally recognised in the thouarsense chronozone (e.g. dean et al. 1961; howarth 1980). the lower or striatulum subchronozone corresponds broadly to the combined bingmanni, striatulum and fascigerum subchronozones of the scheme used here. the upper or fallaciosum subchronozone is probably more or less equivalent to the division of the same name used here. bingmanni subchronozone index. pseudogrammoceras bingmanni (denckman 1887). author. gabilly et al. (1971) as a zone, gabilly (1976) as a subzone. correlating fauna. p. bingmanni and p. struckmanni with rarer podagrosites, etc. included subdivisions. bingmanni zonule (gabilly 1976; elmi et al. 1991, 1994, 1997). striatulum subchronozone index. grammoceras striatulum (j. de c. sowerby 1823). author. buckman (1888 in: buckman 1887–1907). correlating fauna. grammoceras spp. including the index and g. thouarsense with some pseudogrammoceras, etc. included subdivisions. doerntense zonule, thouarsense zonule (gabilly 1976; elmi et al. 1991, 1994, 1997). comment. use of g. striatulum as an index by buckman (1888 in: buckman 1887–1907) has priority over the use of g. thouarsense (used first as a zone by brasil 1896) and also by elmi et al. (1997). fascigerum subchronozone index. esericeras fascigerum (buckman 1888 in: buckman 1887–1907). author. gabilly et al. (1971) as esericeras subzone, guex (1975) as fascigerum subzone. correlating fauna. e. fascigerum and e. spp., pseudogrammoceras differens (ernst) and grammoceras penestriatulum (buckman), etc. included subdivisions. fascigerum zonule (gabilly 1976; elmi et al. 1991, 1994, 1997). fallaciosum (i) subchronozone index. pseudogrammoceras fallaciosum (bayle 1878). author. nicklès (1907), probably as a broader division; replaces the struckmanni subzone of w.t. dean (in: dean et al. 1961) due to the latter’s stratigraphically incorrectly placed index (howarth 1980). correlating fauna. p. grp fallaciosum, osperlioceras, etc. included subdivisions. fallaciosum zonule (gabilly 1976; elmi et al. 1991, 1994, 1997). comment. as a fallaciosum subchronozone is also used in mediterranean province areas, but may not be exactly time equivalent, the northwest european unit is here designated ‘fallaciosum (i)’. dispansum chronozone index. phlyseogrammoceras dispansum (lycett 1860). author. buckman (1889) as dispansum beds, buckman (1910) as dispansum zone. comment. the dispansum chronozone as used here corresponds to the dispansum subzone of the levesquei zone of british authors (e.g. w.t. dean in: dean et al. 1961; howarth 1980). insigne subchronozone index. hammatoceras insigne (zieten 1830). author. welsch (1897) as a zone, theobald & mauberge (1949) as a subzone. correlating fauna. h. ex grp insigne, phlyseogrammoceras dispansum, etc. included subdivisions. cappucinum zonule, pachu zonule (gabilly 1976; elmi et al. 1991, 1994, 1997). gruneri subchronozone index. gruneria gruneri (dumortier 1874). author. gabilly et al. (1971) as a ‘horizon’ (= zonule), used as a subzone by elmi et al. (1991, 1994). correlating fauna. gruneria ex grp gruneri, hammatoceras perplanum prinz, etc. included subdivisions. gruneri zonule (gabilly 1976; elmi et al. 1991, 1994, 1997). 50 51 pseudoradiosa chronozone index. dumortieria pseudoradiosa (branco 1879). author. haug (1892). levesquei subchronozone index. dumortieria levesquei (d’orbigny 1844). author. benecke (1901) as a broader ‘levesqueischichten’, buckman (1925 in: buckman 1909–1930) as a hemera, later as a subzone (spath 1942). correlating fauna. d. ex grp levesquei, catulloceras. included subdivisions. insignisimilis zonule, munieri zonule (gabilly 1976; elmi et al. 1991, 1994, 1997). comment. broadly equivalent to the levesquei subzone of the levesquei zone of british authors (e.g. w.t. dean in: dean et al. 1961; howarth 1980). pseudoradiosa subchronozone index. as pseudoradiosa chronozone (see above). author. used as a subzone by gabilly et al. (1971). correlating fauna. d. ex grp pseudoradiosa including d. explanata buckman, huddlestonia, etc. included subdivisions. pseudoradiosa zonule (gabilly 1976; elmi et al. 1991, 1994, 1997). comments. broadly equivalent to the moorei subzone of w.t. dean (in: dean et al. 1961) although d. moorei was later found to co-occur with the first pleydellia and therefore placed in the aalensis zone, mactra subzone by gabilly (1976). aalensis chronozone index. pleydellia aalensis (zieten 1832). author. reynès (1868). comment. broadly equivalent to the aalensis subzone of w.t. dean (in: dean et al. 1961). mactra subchronozone index. pleydellia mactra (dumortier 1874). author. mouterde (1952) informally as a ‘horizon’, elmi (1967) as a horizon, as a subzone by gabilly et al. (1971). correlating fauna. p. ex grp mactra and the last dumortieria ex grp moorei (lycett) with rarer pseudammatoceras, etc. included subdivisions. tectiforme zonule, mactra zonule, celtica zonule (gabilly 1976; elmi et al. 1991, 1994, 1997). fluitans subchronozone index. pleydellia fluitans (dumortier 1874). author. elmi et al. (1991, 1994). correlating fauna. p. fluitan, p. lugdunensis (dumortier), p. buckmani mauberge, rarer pseudammatoceras, etc. included subdivisions. lugdenensis zonule, buckmani zonule (gabilly 1976; elmi et al. 1991, 1994, 1997). comments. the subzone was created by elmi et al. (1991) as a replacement for the aalensis subzone as p. aalensis itself was considered to be a primitive form of the mactra subzone. the buckmani subzone of gabilly (1976) was also reduced to the status of a ‘horizon’ (= zonule). mediterranean province (italy,austria, southern spain, north africa, etc.) the closest correlations between mediterranean and submediterranean areas are in the bifrons and aalensis chronozones. at other levels, the degree of faunal differentiation has necessitated the creation of the zonal schemes of donovan (1958), elmi et al. (1974), guex (1973) and other authors. a number of horizons (probably mainly with the status of zonules) are tabulated by elmi et al. (1991, 1994) and summarised again by elmi et al. (1997), as reproduced here on figures 6 and 7. lower toarcian substage polymorphum chronozone index. dactylioceras polymorphum fucini 1919. author. jimenez & rivas (1979). equivalent to the mirabile zone of guex (1973; index: d. mirabile fucini). mirabile subchronozone index. dactylioceras mirabile fucini 1919. author. colo (1961) as a ‘horizon’ or ‘niveau’, used as a subzone by guex (1973). correlating fauna. dactylioceras (eodactylites) spp., protogrammoceras (= paltarpites), lioceratoides, etc. included subdivisions. ‘paltarpites’ fauna (?part; elmi et al. 1991, 1994, 1997). semicelatum (ii) subchronozone index and author. as subboreal province (see above), but used in the sense of the submediterranean province. broadly equivalent to the ‘niveau à protogrammoceras madagascariense’ of colo (1961) and the madagascariense subzone of guex (1973; index: protogrammoceras madagascariense (thevenin 1908)). correlating fauna. dactylioceras spp. including d. semicelatum, also protogrammoceras madagascariense thevenin, etc. included subdivisions. ‘paltarpites’ fauna (?part), ‘striatus’ fauna (?part; elmi et al. 1991, 1994, 1997). levisoni chronozone index. hildaites levisoni (simpson 1843). author. guex (1973). levisoni subchronozone index. as levisoni chronozone (see above). author. guex (1973). correlating fauna. hildaites spp. including h. levisoni, also dactylioceras spp., etc. included subdivisions. the ‘striatus’ fauna (?part; elmi et al. 1991, 1994, 1997). falciferum subchronozone index and author. as subboreal province (see above). correlating fauna. harpoceras grp falciferum, mercaticeras, etc. included subdivisions. none (elmi et al. 1991, 1994, 1997). bifrons chronozone index and author. as subboreal province (see above). sublevisoni subchronozone index and author. as submediterranean province (see above). correlating fauna. hildoceras grp sublevisoni, etc. included subdivisions. sublevisoni zonule, tethysi zonule (elmi et al. 1991, 1994, 1997). bifrons subchronozone index and author. as submediterranean province (see above). correlating fauna. hildoceras grp bifrons. included subdivisions. lusitanicum zonule, apertum zonule, bifrons zonule, semipolitum zonule (elmi et al. 1991, 1994, 1997). comment. the semipolitum zonule has been previously used as a full subchronozone (see note under submediterranean province above). upper toarcian substage gradata chronozone index. brodieia gradata (merla 1932). author. atrops & elmi (1971) as a subzone. gemma subchronozone index. gabyllites [collina] gemma (bonarelli 1893). author. gallitelli-wendt (1970). correlating fauna. peronoceras, podagrosites, pseudopolyplectus, etc. included subdivisions. none (elmi et al. 1991, 1994, 1997). alticarinatus subchronozone index. merlaites alticarinatus (merla 1932). author. mouterde (1967) as a ‘niveau’, atrops & elmi (1971) as a subzone. correlating fauna. crassiceras spp., phymatoceras, pseudogrammoceras, polyplectus, etc. included subdivisions. aratum ?zonule, subregale ?zonule (elmi et al. 1991, 1994, 1997). bonarellii chronozone index. hammatoceras bonarellii parisch & viale 1906. author. elmi et al. (1991, 1994) as a replacement, in part, for the rivierense zone of elmi (1986; index osperlioceras rivierense (monestieri)). mediterraneum subchronozone index. pseudogrammoceras mediterraneum rivas 1975. author. used as a subzone by elmi et al. (1991, 1994). correlating fauna. pseudogrammoceras sp. including p. cf. bingmanni, also polyplectus, oxyparoniceras, etc. included subdivisions. none (elmi et al. 1991, 1994, 1997). fallaciosum subchronozone index and author. as northwest european province (see above). correlating fauna. pseudogrammoceras fallaciosum, polyplectus, oxyparoniceras, erycites, etc. included subdivisions. none (elmi et al. 1991, 1994, 1997). speciosum chronozone index. hammatoceras speciosum jamensch 1902. author. used as a zone by elmi et al. (1991, 1994). speciosum subchronozone index. as speciosum zone (see above). author. used as a subzone by elmi et al. (1991, 1994). correlating fauna. hammatoceras insigne, osperlioceras, pseudogrammoceras pachu. included subdivisions. none (elmi et al. 1991, 1994, 1997). reynesi subchronozone index. osperlioceras reynesi (monestier 1921). 52 author. j. mattei (in: gabilly et al. 1971) as ‘assises à pseudogrammoceras reynesi’. correlating fauna. o. reynesi and spp., hammatoceras perplanum, etc. included subdivisions. none (elmi et al. 1991, 1994, 1997). meneghini chronozone index. dumortieria meneghinii haug 1887. author. donovan (1958). correlating fauna. d. meneghinii with catulloceras grp perroudi, osperlioceras, hammatoceras, erycites etc. included subdivisions. meneghinii ?zonule, unnamed interval (elmi et al. 1991, 1994, 1997). aalensis chronozone index and author. as northwest european province (see above). mactra subchronozone index and author. as northwest european province (see above). correlating fauna. pleydellia grp mactra etc. included subdivisions. mactra zonule, flamandi zonule (elmi et al. 1991, 1994, 1997). fluitans subchronozone index and author. as northwest european province (see above). correlating fauna. pleydellia spp. including p. fluitans. included subdivisions. fluitans zonule, unnamed interval (elmi et al. 1974, 1991, 1994, 1997). references alméras, y., boullier, a. & laurin, b. 1997: brachiopodes. in: cariou, e. & hantzpergue, p. 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revision accepted 31 july 1997. geological survey of denmark and greenland bulletin 31, 2014, 31-34 31 calcareous nannofossil and foraminifer biostratigraphy of the campanian–maastrichtian chalk of the femern bælt (denmark–germany) emma sheldon, caterina morigi and sarah d. møller a new study based on calcareous nannofossil and benthic and planktonic foraminifer biostratigraphy is presented for the upper campanian – maastrichtian chalk of the femern bælt (denmark and germany; fig.1). the results are consistent with recent studies of the danish chalk for this interval, allowing correlation across the danish basin and forming the basis for correlation further afield within the boreal realm. numerous studies have been carried out recently on the upper campanian – maastrichtian chalk of the danish basin, covering aspects such as sedimentology, depositional environment, macrofossil biostratigraphy, carbon isotope stratigraphy as well as nannofossil and dinoflagellate biostratigraphy. however, very few published studies on foraminifers exist across this interval in this area. the 09.a.006, 09.a.007 and 09.a.008 boreholes (fig. 2) were drilled in 2009 in preparation for construction of a fixed link across the femern bælt, which will connect denmark to germany (rambøll arup jv 2011). the boreholes penetrated glacial till, paleocene–eocene clay and chalk (sheldon et al. 2012). here, for the first time, the boreal foraminifer biostratigraphy of the late campanian – maastrichtian interval is investigated and presented alongside nannofossil biostratigraphy. geological setting and palaeogeography the femern bælt area is located to the south of the ringkøbing–fyn high and is part of the german basin (fig. 1). during the late cretaceous the danish area was part of the extensive epicontinental sea where cool-water carbonate deposition dominated. the maastrichtian chalk was deposited at depths of 100–250 m (surlyk 1997). chalk distribution patterns are affected in the danish area by folding, salt diapirism, non-deposition and erosion, especially in the southern region (lieberkind et al. 1982). in the maastrichtian to danian, the area was situated at 44°–46°n (smith et al. 1994). in the upper maastrichtian of this part of the danish area, two main facies types dominate: (1) relatively deep water basinal sediments comprising coccolith and foraminferarich pelagic chalk, and (2) shallow marine chalk with high diversity faunas dominated by bryozoans, echinoids, bivalves and brachiopods (surlyk 1997; hart et al. 2004). the upper campanian – maastrichtian chalk of onshore denmark was until recently referred to the tor formation equivalent. the tor formation was established by deegan & scull (1977) as a maastrichtian (locally upper campanian) chalk unit in the norwegian and danish sectors of the north sea. a new holostratigraphic analysis of the upper cretaceous chalk of eastern denmark resulted in a lithostratigraphic subdivision of the onshore chalk (surlyk et al. 2013). however due to local facies variations this new subdivision cannot be applied to the chalk of the femern belt area with© 2014 geus. geological survey of denmark and greenland bulletin 31, 31–34. open access: www.geus.dk/publications/bull fehmarn south north lolland09.a.006 09.a.00809.a.007 3 km 0 m 50 100 150 quaternary deposits folded and faulted palaeogene clay palaeogene cretaceous fig.1. map of denmark and northern germany showing the location of the planned fixed road and rail link across femern bælt. l: lolland, f: fehmarn, fb: femern bælt. fig. 2. sketch south–north cross-section of the femern bælt area and the location of the boreholes (from sheldon et al. 2012). 55° 10°e 100 km 14°e 57° sweden denmark germany german basin danish basin norway l f fb 54° 56°n 58° poland ringkøbing–fyn high 3232 out sedimentological analysis and the term ‘tor formation equivalent’ is retained here. biostratigraphy from the 09.a.008 and 09.a.006 boreholes nannofossils and foraminifers from the campanian–maastrichtian were analysed; from the 09.a.007 borehole nannofossils from the upper maastrichtian were examined. the southern north sea foraminifer zonation of king et al. (1989) and the boreal nannofossil scheme of burnett (1998) were applied (fig. 3). marker species are shown in fig. 4 and biostratigraphic results on fig. 5. nannofossil zones – the co-occurrence of orastrum campanensis and eiffelithus eximius indicates the presence of uc15dbp at the base of core 09.a.008. subzone uc15ebp is absent. uc16bp was defined by burnett (1998) but recent studies of danish campanian–maastrichtian chalks (sheldon 2008; thibault et al. 2012) question the reliability of the uc16bp marker species in the danish area. the top of uc16abp is defined by the last occurrence (lo) of heteromarginatus bugensis. in the present study and in the stevns-1 borehole, eastern denmark, the lo of h. bugensis is below the lo of tortolithus caistorensis (marker for top uc16bbp). other studies (e.g. fritsen et al. 1999) assign the lo of h. bugensis to the top of uc16cbp. the top of uc16bbp is defined by the lo of t. caistorensis, which is very rare in the danish area. uc16abp and b are not easily subdivided in the danish area and alternative markers are suggested for the top of uc16bbp: the lo’s of tortolithus hallii and tortolithus pagei were successfully applied in the present study, positioned prior to the los of broinsonia parca parca and zeugrhabdotus praesigmoides (markers for the overlying uc16cbp). the top of uc16cbp is defined by the lo of monomarginatus quaternarius. the lo of b. parca parca as a top uc16cbp marker is a more reliable marker in this study. the lo of z. praesigmoides as an additional top uc16cbp marker is confirmed in this study. the top of subzone uc16d bp is defined by the lo of broinsonia parca constricta. the first occurrence (fo) of prediscosphaera mgayae as an additional marker c am pa ni an m aa st ric ht ian southern north sea foraminifer zonation boreal calcareous nannofossils fc s2 3 fc s2 2 la te c re ta ce ou s fc s2 1b a a b c a b c u c 16 u c 20 u c 15 d d b. draco b. miliaris n. frequens a. maastrichtiana c. daniae cretaceous taxa l. quadratus r. levis t. orionatus b. parca constricta m. quaternarius t. caistorensis h. bugensis e. eximius a. bettenstaedti b. decoratus g. monterelensis b palaeogene epoch/ageage (ma) uc19 uc18 uc17 70 75 65 72.1 e m lt lt lt 66.0 a b c d e f g h i j k l m n o p q r s t u v w fig. 3. nannofossil and foraminifer zonations. the timescale is according to gradstein et al. (2012). ucbp: upper cretaceous boreal province nannofossil zones of burnett (1998), fcs: cretaceous southern north sea foraminfers, shelf facies, including chalk, zones of king et al. (1989). nannofossil and foraminifer zonations correlated using fritsen (1999). fig. 4. sselected nannofossils and foraminifers from the femern bælt. a: arkhangelskiella maastrichtiana. b: prediscosphaera  stoveri. c: nephrolithus  frequens. d: reinhardtites levis. e: tranolithus orionatus. f: prediscosphaera mgayae. g: broinsonia  parca constricta. h: monomarginatus  quaternarius. i: zeugrhabdotus praesigmoides. j: tortolithus caistorensis. k: heteromarginatus bugensis. l: orastrum campanensis. m: bolivinoides draco. n: pseudouvigerina cristata. o: brizalina incrassata. p: hagenowella paleocenica. q: bolivinoides draco giganteus. r: bolivinoides decoratus. s: stensioeina pommerana. t: angulogavelinella bettenstaedti. u: globotruncana arca. v: gavelinella monterelensis. w: globorotalites micheliana. scale bars, nannofossils: 5 µm, foraminifers: 0.1 mm. 33 in the middle of uc16d bp (thibaut et al. 2012) is confirmed in this study. the lo of tranolithus orionatus marks the top of uc17 bp. in the present study, the lo of this species occurred after the lo of reinhardtites levis (the lo of which marks the top of the overlying uc18 bp). in stevns-1 these species also occur in the ‘reverse’ order. the lo of r. levis marks the top of uc18bp, but in this study its lo is before that of t. orionatus. it is suggested that in the danish area, uc17bp and uc18bp should be combined, using either the lo of t. orionatus or the lo of r. levis to mark the zone top. the lo of p. mgayae has been suggested as an additional marker for the top of this zone (thibaut et al. 2012); this is confirmed in this study. uc19bp comprises the interval from the lo of reinhardtites levis to the fo of lithraphidites quadratus. the fo of l. quadratus defines the base of uc20abp. the base of uc20bbp is defined by the fo of nephrolithus frequens. the base of uc20cbp is defined by the fo of arkhangelskiella maastrichtiana. the uncertainties surrounding the use of the fo of a. maastrichtiana as a marker are well-documented (e.g. thibault 2010) and the two subzones are merged here. the first common occurrence of a. maastrichtiana is used to mark the base of uc20b-cbp in this study. the base of the prediscosphaera stoveri acme within this combined subzone is also applied successfully in this study. the lo of helicolithus trabeculatus was suggested as a supplementary marker for this level (sheldon 2008) and is confirmed here. the fo of cribrosphaerella daniae marks the base of uc20dbp, the uppermost subzone of the maastrichtian. the co-occurrence of chiasmolithus edentulus, coccolithus pelagicus, cyclagelosphaera alta and neochiastozygus saepes assigns the base of the overlying danian in borehole 09.a.008 to the upper danian subzone nntp4f (varol 1998). lower – upper danian subzones nntp1a-4f are missing or were not sampled. danian chalk in the femern area was documented for the first time recently (sheldon et al. 2012). previously the southern limits of danian deposits were thought to be farther to the north (håkansson & pedersen 1992). foraminifer zones – the top of fcs23 is defined by the lo of pseudotextularia elegans. bolivinoides draco (s.s.) is an index species. the top of subzone fcs23a is defined by the fo of p. elegans. p. elegans is not seen in the present study, but rare occurrences of bolivinoides draco were noted, indicating the fig. 5. nannofossil and foraminifer biostratigraphic correlations of the three boreholes at the femern bælt. 20 25 30 35 40 45 50 55 60 65 70 75 80 85 22 ac 21 b 23 90 95 100 20 bc 20 a present 09.a.006 09.a.007 09.a.008 fo lo foraminifersf n common a. maastrichtiana common a. maastrichtiana c. daniae and danian flora l. quadratus t. orionatus, r. levis p. mgayae p. mgayae b. parca constricta z. praesigmoides tortolithus spp. h. bugensis o. campanensis, e. eximius n. frequens, common a. maastrichtiana p. cristata total depth 100.6 m total depth 50.1 m total depth 99.9 m b. draco giganteus b. draco, spirillina spp. a. bettenstaedti g. arca barren p. cristata s. pommerana h. paleocenica (black) h. paleocenica (black) n. frequens n. frequens l. quadratus l. quadratus t. orionatus, p. mgayae b. parca constricta r. levis p. mgayae m. quaternarius tortolithus spp. z. praesigmoides h. bugensis 20 bc 20 a 19 17/18 16 d 16 c 16 b 23 fc s 22 ab 21 -b 20 bc 20 d 20 a 19 17 -1 8 16 c 16 b 16 a 15 d 16 d ucfcs f n f nn uc ucfcs depth (m) b. draco, h. paleocenica (black) h. paleocenica (black) p. cristata spirillina spp. b. draco giganteus,s. pommerana p. cristata barren g. arca, ? r. szajnochae g. monterelensis g. arca globotruncana spp., a. bettenstaedti b. miliaris, a. bettenstaedti b. draco giganteus nannofossils (first occurrence) (last occurrence) 3434 presence of fcs23. in upper fcs23 the range of hagenowella paleocenica and the fo of spirillina spp. are additional markers in the present study. the lo of stensioeina pommerana occurring towards the bottom of fcs23 (corresponding to the base of nannofossil subzone uc20abp) in this study may also prove useful. the range of pseudouvigernina cristata within fcs23 may prove to be another useful marker. the absence of p. elegans means that it is not possible to subdivide fcs23 in this study. the top of fcs22 is defined by the lo of bolivinoides miliaris. fcs22 is divided into fcs22b, the top of which is defined by the lo of b. miliaris, and fcs22a, the top of which is defined by the lo of angulogavelinella bettenstaedti. the los of a. bettenstaedti and b. miliaris occurred in the same sample suggesting the presence of only fcs22a, but the range of a. bettenstaedti within fcs22 could be useful. the top of fcs21 is defined by the lo (often a flood occurrence) of reussella szajnochae. fcs21 is divided into fcs21b, the top of which is defined by the lo of r. szajnochae, and fcs21a, the top of which is defined by the lo of gavelinella usakensis. g. usakensis was not found in the present study, indicating the presence only of fcs21b. the lo of globotruncana arca at the top of fcs21 is an additional marker, as are the los of gavelinella monterelensis and globorotalites micheliana and fo of brizalina incrassata at a slightly lower stratigraphic level. conclusions the campanian–maastrichtian chalk of the danish area has recently been studied intensively using calcareous nannofossils, applying the ucbp scheme. the studies underline the need for an amendment of the uc scheme for the danish area. the present study highlights the need to reconsider the subdivision of uc20bbp and uc20cbp based on the fo of arkhangelskiella maastrichtiana. it is recommended for the danish area to use the fo of common a. maastrichtiana to mark the base of uc20cbp. it may also be practical to merge uc17bp and uc18bp. additionally the subdivision of uc16bp cannot be reliably applied in the danish area. the danian nannofossil assemblage in borehole 09.a.008 provides further evidence for the southerly encroachment of the danian sea. the foraminifer zonation of the late campanian – maastrichtian is for the first time correlated with the nannofossil zonation based on the 09.a.006 and 09.a.008 cores (fig. 5). in the absence of certain established fcs marker foraminifera for the north sea chalk, new zonal markers, e.g. hagenowella paleocenica, pseudouvigerina cristata, globotruncana arca, gavelinella monterelensis and globorotalites micheliana are used here in addition to conventional taxa for the campanian–maastrichtian chalk of denmark, perhaps allowing correlation with the north sea area, and further afield. references burnett, j.a. 1998: upper cretaceous. in: bown, p.r. (ed.): calcareous nannofossil biostratigraphy. british micropalaeontological society series 5, 132–199. deegan, c.e. & scull, b.j. 1977: a standard lithostratigraphical nomenclature for the central and northern north sea. the institute of geological sciences report 77/25, 36 pp. fritsen, a. (ed.) 1999: a joint chalk stratigraphic framework. in: joint chalk research program topic v 1. norwegian petroleum directorate. 206 pp. gradstein, f.m., ogg, j.g., schmitz, m.d. & ogg, g.m. 2012: the geologic time scale 2012, 1176 pp. amsterdam: elsevier. hart, m.b., feist, s.e., price, g.d. & leng, m.j. 2004: reappraisal of the k–t boundary succession at stevns klint, denmark. journal of the geological society (london) 161, 885–892. håkansson, e., & pedersen, s.a.s. 1992: geologisk kort over den danske undergrund. map sheet. københavn: varv. king, c., bailey, h.w., burton, c.a. & king, a.d. 1989: cretaceous of the north sea. in: jenkins, d.g. & murray, j.w. (eds): stratigraphic atlas of fossil foraminifera, 372–417. chichester: ellis horwood. lieberkind, k., bang, i., mikkelsen, n. & nygaard, e. 1982: late cretaceous and danian limestone. danmarks geologiske undersøgelse serie b 8, 49–62. rambøll arup jv 2011: summary of geological conditions. geotechnical data report 01.3-002, 53 pp. virum: femern a/s. sheldon, e. 2008: upper campanian–maastrichtian calcareous nannofossil biostratigraphy of the stevns-1 borehole, denmark. journal of nannoplankton research 30, 39–49. sheldon, e., gravesen, p. & nøhr-hansen, h. 2012: geology of the femern bælt area between denmark and germany. geological survey of denmark and greenland bulletin 26, 13–16. smith a.g., smith, d.g. & funnel, b.m. 1994: atlas of mesozoic and cenozoic coastlines, 99 pp. cambridge: cambridge university press. surlyk, f. 1997: a cool-water carbonate ramp with bryozoan mounds: late cretaceous – danian of the danish basin. sepm special publication 56, 293–307. surlyk, f., rasmussen, s.l., boussaha, m., schiøler, p., schovsbo, n.h., sheldon, e., stemmerik, l. & thibault, n. 2013: upper campanian − maastrichtian holostratigraphy of the eastern danish basin. cretaceous research 46, 232–256. thibault, n. 2010: biometric analysis of the arkhangelskiella group in the upper campanian–maastrichtian of the stevns-1 borehole, denmark: taxonomic implications and evolutionary trends. geobios 43, 639–652. thibault, n., harlou, r., schovsbo, n., schiøler, p., minoletti, f., galbrun, b., lauridsen, b.w., sheldon, e., stemmerik, l. & surlyk, f. 2012: upper campanian – maastrichtian nannofossil biostratigraphy and high resolution carbon isotope stratigraphy of the danish basin: towards a standard δ13c curve for the boreal realm. cretaceous research 33, 72–90. varol, o. 1998: palaeogene. in: bown, p.r. (ed.): calcareous nannofossil biostratigraphy. british micropalaeontological society publication series, 200–224. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: es@geus.dk http://store.elsevier.com/the-geologic-time-scale-2012-2-volume-set/isbn-9780444594259/ http://store.elsevier.com/the-geologic-time-scale-2012-2-volume-set/isbn-9780444594259/ mailto:obe@geus.dk geological survey of denmark and greenland bulletin 38, 2017, 65-68 65 in 2009, the danish energy agency informed the geological survey of denmark and greenland (geus) that a new legal act (lov om fordring/forældelsesloven) would become effective in 2011. the new act introduced a limitation period of three years on the government’s right to request data from licensees’ oil and gas activities in denmark (www.retsinformation.dk 2007), and it became the catalyst for a major change in the standard procedure for licensees’ submission of seismic data related to exploration activities to geus. a consequence of the new legal act was that the danish authorities would have to request data from the licensee (following the consolidated act of the use of the subsoil) from the licensee no later than three years after their generation or publication; otherwise the authorities would have no legal right to the data. it was emphasised by the danish energy agency that geus would have to submit all requests for outstanding data to the licensees no later than 1 january 2011. during the 1980s and 1990s the standard procedure for data requests from geus did not include pre-stack seismic data, e.g. raw field data. a search in geus database showed that field data had been received from only a small percentage of existing 2d and 3d surveys (fig. 1). being the national data bank for geological data, geus had to respond to this risk of losing access to valuable data, and immediate action was needed. the national well data and subsurface archives geus and its danish predecessor the geological survey of denmark (dgu) have systematically collected and stored geological data from denmark for more than a 100 years. the idea of establishing a national data bank for geological data arose in the 1920s when danish law made it mandatory for everyone drilling in search for water to register certain geological information encountered during drilling and deliver it to dgu. the well data archive was established in 1926 to fulfil this purpose (troelstrup 1992). the search for oil in denmark began onshore in 1935. until the beginning of the 1980s data submitted to dgu from oil and gas activities were stored in the well data archive. for many years only few companies searched for oil in the danish sector. in 1962, the danish company a.p. møller a/s was granted exclusive rights to explore the danish subsoil according to the sole concession of 8 july 1962 (north the rescue of seismic field data from exploration activities in the danish north sea marianne m. hansen and nicolai rinds 5°e 3°e 6°e 55°n 56°n norway denmark 500 km 50 km fig. 1. 3d seismic field data status prior to the rescue project. green polygons: field data in geus’ archive. red polygons: field data not in geus’ archive. grey polygons: parts of seismic surveys outside the danish sector. 30 000 20 000 2d km 3d km2 10 000 0 1966 1970 1974 1978 1982 1986 1990 1994 1998 2002 2006 2010 2014 fig. 2. annual seismic survey activities in denmark 1962–2014. © 2017 geus. geological survey of denmark and greenland bulletin 38, 65–68. open access: www.geus.dk/publications/bull 6666 sea fund 2012). this was renegotiated in 1981, and a large concession area was relinquished. in 1984 the first competitive licensing round resulted in an increased search for oil, and the submission of data to dgu increased (fig. 2, extracted from geus’ database). in the beginning of the 1980s, it was decided that a separate archive should be established – the subsurface archive – to store oiland gas-related data (kristoffersen 1995, 2017). the decision was made in order to fulfil increased requirements for a secure data environment and to control the strict confidentiality regulations applying to exploration and production data. data usage until 1980, dgu did not request digital seismic data from, the companies as part of its standard procedure. this was primarily due to the fact that only paper copies of seismic sections were used for in-house interpretation. furthermore, space for in-house storage was very limited, and the data owners themselves were obliged to store their digital field data for a period of time and to offer the data to dgu if they decided to discard them. at that time the data were stored on vast numbers of 9-track tapes. in general, the oil companies did not use digital data either, and if such data were requested it was often on a single line basis. by 1990, workstations had become a standard tool for interpretation and geophysical mapping in the oil industry, and access to the digital data became necessary. therefore digital processed seismic data were added to dgu’s standard data request, but their handling was a huge and timeconsuming task. licensees or companies were asked only to submit specific digital seismic data when they were needed for in-house interpretation, or if a specific request arrived from an external customer. as a consequence, dgu did not have a complete collection of digital processed seismic data at the time. until the 1990s, most of the seismic field data submitted to dgu came from relinquished or expired licences. in these cases, the data holders could decide to keep the seismic field data themselves or offer the data to dgu. although the authorities have always been able to request copies of the seismic field data this was not listed as a mandatory requirement until 2002 (danish energy agency 2002). data media and volumes the development of 3d data acquisition methods in the late 1980s resulted in an increasing amount of tapes being submitted to dgu. during the last 20 years, digital seismic data have regularly been submitted to the subsurface archive and maintenance of these data has been a constant challenge. the long-term use of the digital data can only be secured by keeping the data from degrading, by rewinding the 9-track tapes or transcribing the data to a new media on a regular basis. the oil industry has frequently changed their ‘standard tape media’, in order to secure the very expensive collected data and to reduce the storage space needed on board vessels and in archives. the most commonly used storage medium in the 1980s and 1990s was 9-track tapes. in the mid-1990s, more compact types appeared such as exabyte, dat and various versions of ibm tapes, which were smaller and could contain much more data per volume. this had a tremendous impact on the number of shelf metres needed for storage. for instance, one modern ibm3592-jd tape contains the same amount of data as 70 000 high-capacity 9-track tapes, which would require 1750 archive shelf metres (fig. 3). when transcribed storage space equals 6 lorries input: 70 000 9-track tapes of 145 mb (1750 shelf metres) output: 1 tape ibm3592-jd of 10 tb fig. 3. space reduction with the use of modern tapes media. source: ovationdata 2017. fig. 4. seismic data and tape drives at geus. 67 over time, geus has purchased various kinds of tape drives to be able to read the data and (on a small scale) make backup copies. in 2000, the subsurface archive contained data stored on up to 10 different tape media types, but mostly on c. 15 000 9-track tapes, including both seismic and well data, as well as data acquired in greenland waters. the archive also contained 3000 exabyte, dat and lto tapes and 5000 ibm tapes of various kinds. at geus tape copying is considered a specialist job since up-to-date equipment and specialised treatment of deteriorating 9-track tapes are required, and the task is very time consuming, expensive and never-ending. furthermore, vintage tape drives are difficult to maintain and acquisition of spare parts almost impossible. since 2002, geus has used an external company for all tape copying jobs. a few in-house tape drives are still being used and maintained for internal purposes (fig. 4). the data rescue project after geus was made aware of the new act and its consequences, it was decided in 2010 to start a project to rescue older seismic field data. the aim was to rescue i.e. locate, receive and store either the original or a copy of the seismic data from all oiland gas-related activities in denmark. the plan was within the first year to contact all companies which had been operators or had acquired data in the danish sector between 1980 and 1999. the plan for the following two years was to perform quality control of the submitted data, with completion of the project within three years. the following priorities were set up: field data were considered more important than processed data. 3d data were considered more valuable than 2d data. finally, onshore 2d data were considered more important than offshore 2d data, and data acquired from 1990 to 2000 were considered to be more valuable than older data. a search in geus’ database showed that almost no field data from before 1995 had been received, and that some field data had been received between 1995 and 2002. worst-case scenario a worst-case scenario in terms of expenditure and workload for geus was considered, assuming that all the companies chose to give up all their original field tapes and submit them to geus. it might be expected from the age distribution that most data might still be stored on the original 9-track tapes. if geus had to accept these directly, the subsurface archive would need large additional resources to cover the costs of external storage or for transcription to an in-house, modern storage medium. a rough calculation showed that if all the missing field data were submitted, geus could receive more than 190 000 tapes, out of which 170 000 were likely to be 9-track tapes. the cost of transcription was estimated to c. dkk 23 million (table 1). alternatively, if all the 170 000 9-track tapes were to be stored in-house the subsurface archive would need 4350 additional shelf metres. after making a list of 2d and 3d surveys from which the seismic field data had not been submitted, a priority list consisting of 48 3d surveys and 40 2d surveys was made. the next step was to contact the companies involved, but it proved difficult to locate some of the companies and relevant contact persons. in 2002, geus had sent requests for data to all licensees in the danish sector, typically to the company headquarters without a specific contact person. the outcome of this campaign was unfortunately almost 2d 3405 3405 140 476 700 3d 156 234 156 234 140 21 872 760 total cost 22 349 460 †assuming: all field data from 1995 and earlier are located. all data are stored on 9-track tapes. data are 60-fold. there is 1 km of data per tape. survey type total km in no. of tapes dkk per tape dkk total denmark table 1. estimated copying cost – worst-case scenario† 4°e3°e 5°e 6°e 55°n 56°n 50 km fig. 5. 3d field data status after the rescue project. for location and polygon colours, please see fig. 1. 6868 nill. the oil industry is a rapidly changing business, and some of the companies which had been active in the danish sector in the 1970s or 1980s no longer existed or had been taken over by or merged with other companies. keeping track of a company can be strenuous and time consuming, and the subsequent search for a relevant contact person was even more difficult. results and achievements as part of the data rescue project, 17 companies were contacted during 2010 and 2011. in general their responses to locate and deliver the missing data have been positive and cooperative. the project benefitted from the fact that two of the older oil companies are still operating in the danish sector, since both mærsk olie og gas a/s (mærsk) and dong energy e&p a/s (dong) have been in charge of, or involved in, data acquisition for many years. the most time-consuming part of this project turned out to be the quality control of the received data due to insufficient documentation, e.g. erroneous transmittals or missing acquisition and processing reports. the recovery process was initiated in 2010 by meetings with the two companies, whereby a large number of surveys from the priority list could be ticked off, and the focus then shifted towards contacting the owners of the remaining data. mærsk and dong decided to transcribe their original data themselves to modern media types. given the large quantities of data and tapes, this task was very time consuming. it took several years before the copying jobs were finalised and tapes could be submitted to geus, but it had the benefit that geus received all data on modern media types (ibm 3592 tapes or usb disks). initially, high-priority 3d field data from 34 of the 48 3d surveys on geus’ priority list were missing; to date geus has located 82% of the missing surveys and received data from 74% of them (fig. 5). geus has also received some missing processed data. because of the above-mentioned complications, the recovery project lasted six years instead of the anticipated three, with a time consumption of c. 1600 man hours at geus. over a six-year period geus has spent c. 1 million dkk on external copying. lessons learnt and conclusions in hindsight, the timing of the rescue project was optimal. between 2010 and 2014 the world experienced high oil prices (www.macrotrends.net 2017), which is likely to have encouraged the oil companies, especially mærsk and dong, to transcribe their original field tapes to modern media. the companies took the opportunity to recover their vintage data and thereby also save future storage costs. if the project was to be started today, the companies might have had other priorities and it might not have been possible for geus to keep the copying costs at a relatively low level compared to the worst-case scenario. to rescue seismic field data is both a cumbersome and potentially expensive task. however, seismic field data represent a valuable asset for evaluation of the hydrocarbon potential of a given area. new companies in the danish sector, especially smaller companies, commonly request field data from both 2d and 3d seismic surveys in order to reprocess existing data prior to committing themselves to acquisition of new seismic data. the access time to the field data is optimised, since the rescue project has provided geus with most of the field data requested by the companies. the data can also be used in future scientific studies. it is therefore an important task for a geological survey like geus to secure these data by making it a high priority to request and secure all data which are acquired from the danish subsurface. references danish energy agency 2002: danish executive order no. 56 of february 4, 2002 executive order on submission of samples and other information about the danish subsoil. danish executive order no. 56. [unofficial translation]. https://ens.dk/sites/ens.dk/files/oliegas/ submission_samples_other_info_danish_subsoil.pdf kristoffersen, f.n. 1995: dgu’s undergrundsarkiver [dgu’s subsurface archives]. in: binzer, k. (ed.): annual report 1994. copenhagen: geological survey of denmark, 102–105 [in danish]. kristoffersen, f.n. 2017: undergrundsarkiv fra kaos til moderne databank [the subsurface archive – from chaos to modern data bank]. hundested, 1 p. [unpublished report in danish]. north sea fund 2012: brief history of danish oil & gas exploration 1984 – 2011 – licensing rounds 1–6. renewed interest in hydrocarbon exploration in denmark. ahead of the 7th round in 2013, 14 pp. troelstrup, s. 1992: the national geological databank. in: binzer, k. (ed.): annual report 1991. copenhagen: geological survey of denmark, 45–47. www.macrotrends.net 2017: http://www.macrotrends.net/2516/wti-crudeoil-prices-10-year-daily-chart www. retsinformation.dk 2007: lov nr 522 af 06/06/2007 (forældelsesloven) – historisk. https://www.retsinformation.dk/forms/r0710. aspx?id=2655 [in danish – no translation available]. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: mmh@geus.dk. geological survey of denmark and greenland bulletin 31, 2014, 23-26 23© 2014 geus. geological survey of denmark and greenland bulletin 31, 23–26. open access: www.geus.dk/publications/bull seismic stratigraphy and sedimentary architecture of the chalk group in south-west denmark connie larsen, jon ineson and lars ole boldreel the chalk group is ubiquitous in the subsurface of the danish basin and its upper levels are exposed locally onshore, most notably in eastern denmark. although many subsurface studies have been made of the group in the danish basin, most of these have been in the eastern part of denmark (e.g. esmerode et al. 2007; surlyk & lykke-andersen 2007) whereas the stratigraphy and character of the chalk group in the western onshore region is less well-known. the work described here was undertaken as a bsc project at the department of geosciences and natural resource mangement at the university of copenhagen by the first author as part of regional seismic mapping work contributing to an evaluation of the geothermal energy potential in denmark. the aim of this paper is to present a summary of the key results of the study. we have subdivided and mapped the distribution of the chalk group in the northern north german basin and the south-western danish basin based on digital reflection seismic profiles. we also highlight seismic architectural features that testify to periods of active bottom currents. geological setting following rifting in late carboniferous – early permian times, the mesozoic history of the danish basin was dominated by thermal subsidence, albeit with an important midjurassic uplift event and phases of localised inversion in the late cretaceous – palaeogene (liboriussen et al. 1987). the combination of high sea-levels, a peneplaned hinterland and aridity in northern europe in the late cretaceous led to a reduced influx of siliciclastic sediment to the sea and favoured the accumulation of pure pelagic chalk composed primarily of skeletal fragments from coccolithophorid algae. the pelagic chalk deposits were subject to redistribution by various processes including downslope mass-flow movements from 50 km ringkøbing – danish basin tornquist zone fyn high a b jylland north sea 55°n 57°n sorgenfrei– 10°e <500 m 500–1000 m 1000–1500 m 1500–2000 m >2000 m fault salt dome late cretaceous inversion basement high 50 km 10°e seismic line (conventional) seismic line (high resolution) kegnæs-1 boreholedcj-7 (fig. 3) north german basin fig. 1. structural map of the danish basin (a) showing the thickness of the upper cretaceous – danian chalk group (modified after rasmussen & surlyk 2012); note that the study area (shown in detail in b) straddles a major structural high, the ringkøbing–fyn high. the enlarged map (b) displays the seismic lines and deep boreholes used in this study; highlighted in red are the kegnæs-1 borehole (see fig. 2) and the seismic line illustrated in fig. 3. 2424 slope instability caused by syndepositional tectonics as well as along-slope bottom currents that sculpted the sea floor into valleys, channels, drifts, ridges and mounds (e.g. surlyk & lykke-andersen 2007). results the subdivision of the chalk group is based on seismic facies analysis and seismic structural analysis, and the resultant seismic units are correlated to wireline-logs and stratigraphic units from seven deep boreholes (fig. 1). the study used 2d digital and scanned seismic profiles from nine surveys on and offshore; the digital high-resolution and conventional offshore data are of high quality whereas onshore the digital data are of good quality and the few older scanned data are of low quality. the interval between the base and top of the chalk group on the seismic profiles is divided into seismic units based on the recognition and mapping of relatively strong, bounding reflectors. the brief descriptions of the units given below are based on the interpretation of all the seismic profiles in the study area. in the northern part of the study area, the chalk group is subdivided into 10 seismic units but the uppermost unit is absent south of the ringkøbing–fyn high, i.e. only nine units are recognised in this area (figs 1, 2). the inferred ages of these units are based on correlation to dated borehole sections. seismic units 1–3 show parallel, continuous or discontinuous reflector patterns with low amplitude. boundaries of unit 1 are clearly defined on gamma-ray logs (fig. 2) whereas the boundaries of unit 3 are typically marked by minor log fluctuations. units 1–3 are recognised throughout the study area and are broadly of cenomanian, turonian and coniacian ages respectively. seismic units 4–6 display discontinuous, semi-parallel internal reflector patterns which are locally chaotic. boundaries can be correlated with minor gamma-ray log fluctuations in most boreholes; the base of unit 6 is marked by a prominent sonic peak log in the kegnæs-1 borehole (fig. 2). units 4–6 are recognised throughout the study area and are of probable santonian, early campanian and late campanian ages respectively. seismic units 7 and 8 show discontinuous, semi-parallel internal reflector patterns that become chaotic locally. correlation to log data in boreholes can be difficult but a clear response is observed on the gamma-ray logs in some of the boreholes; the unit is recognised throughout the study area and probably of early to mid-maastrichtian age. stratigraphy danian pa le oc en e u pp er c re ta ce ou s up pe r lo w er lo w er c re t. maastrichtian campanian santonian coniacian turonian cenomanian albian seismic unit age ma 9 65.6 70.6 83.5 85.8 89.3 93.5 99.6 8 7 200 300 400 tw ow ay tr av el ti m e (m se c) 500 600 700 6 5 4 3 2 1 500 m kegnæs-1 sonic gamma fig. 2. seismic profile (in two-way travel-time, twt) intersecting the kegnæs-1 borehole, illustrated by the sonic log (blue) and gamma-ray log (red), and showing the correlation between the seismic units and the chronostratigraphy, based on data from gearhart (1986). 25 seismic unit 9 commonly shows localised disturbance of the internal reflector pattern but elsewhere exhibits a parallel, continuous to discontinuous pattern. the unit boundaries are readily correlated to gamma-ray logs and are recognised throughout the study area. unit 9 is of probable late maastrichtian age. seismic unit 10 is only present in the western part of the study area, north of the ringkøbing–fyn high, and wedges out to the south near the ringkøbing–fyn high. it shows a parallel, continuous or discontinuous internal reflector pattern with high amplitude. it is recognisable on some of the gamma-ray logs and is probably of early danian age. based on seismic facies analysis, seismic units 1–3 and 9 (cenomanian, turonian, coniacian and upper maastrichtian ages respectively) typically display parallel and low-angle, mounded internal geometries, whereas seismic units 4–8 (santonian, campanian, lower maastrichtian) show lowangle, mounded, sigmoidal, oblique and hummocky forms (fig. 3). evidence of seismic truncation, linked in some cases to log or biostratigraphic data, indicates the presence of three unconformities in the upper chalk group at ‘mid-campanian’ (base unit 6), ‘base maastrichtian’ (base unit 7) and ‘top maastrichtian’ (top unit 9, base unit 10 where present); these surfaces seem to be associated with complex systems of major drift, minor mounded drift and channel-like features (fig. 3). discussion mapping the distribution and thickness of the chalk group seismic units in the study area illustrates both regional trends, related to active inversion of the sorgenfrei–tornquist zone during deposition of the chalk group, and local trends related to the influence of the ringkøbing–fyn high. the structural control exerted by late cretaceous inversion along the sorgenfrei–tornquist zone is reflected by the over400 600 800 2000 m a b c d e seismic unit 5, probably early campanian in age seismic unit 7, probably early–mid-maastrichtian in age su 5 su 7 su 8 su 9 seismic unit 8, probably early–mid-maastrichtian in age seismic unit 9, probably late maastrichtian in age nw sse dcj-07 nw sse 1000 m nw sse 1000 m nw sse 1000 m 1000 m a a a b d e b c c d e b tw ow ay tr av el ti m e (m se c) fig. 3. internal architecture of selected seismic units in the seismic profile dcj-07 (for location, see fig. 1): unit 5 (green), unit 7 (pink), unit 8 (blue) and unit 9 (purple). the accompanying sketches illustrate the internal architecture of these units; the bounding surfaces indicated in colour are defined on fig. 2 . 2626 all geometry of the chalk group and the gentle northward tilt of the base of the chalk group (see also lykke-andersen & surlyk 2004). the influence of the ringkøbing–fyn high, at least episodically, is indicated by both the distribution of the danian strata and by the evidence of active bottom current activity. the danian chalk succession (seismic unit 10) is limited to the area north of the ringkøbing–fyn high and locally these deposits are preserved in trough-like, channel features. this distribution is consistent with the work of thomsen (1995) who reported that the danian succession wedges out in the south-western part of jylland, crossing the ringkøbing–fyn high. major drifts, minor mounded drifts and channel-like features are identified north of the ringkøbing–fyn high and represent a complex depositional system that is mainly recognised in the campanian–maastrichtian interval; such features were not identified south of the ringkøbing–fyn high. most of the structures are recognised on nw–se and nnw–sse-oriented seismic lines indicating, together with internal architecture, drift migration towards the nw or nnw. a comparable, complex system has been described in the øresund area (esmerode et al. 2007; surlyk & lykkeandersen 2007) where it is seen in middle santonian – campanian and upper maastrichtian intervals, and is considered to be analogous to modern deep-water contourite systems; i.e. that the deposits were moulded by bottom currents that flowed parallel to bathymetric contours. such processes do not seem to have influenced the santonian succession in the study area of south-west denmark, possibly because there was no significant bathymetric relief in this part of the basin at that time. in addition to the upper surface of the chalk group south of the ringkøbing–fyn high, two intra-chalk unconformities are recognised seismically over the entire study area, the inferred mid-campanian hiatus and the ‘base maastrichtian’ surface; note that the former is correlated to a pronounced cemented layer (hardground?) and a biostratigraphically defined middle campanian hiatus in the kegnæs-1 borehole (gearhart 1986). according to abramovitz et al. (2010), two major unconformities are recognised within the chalk group in the southern part of the danish central graben: one within the turonian–campanian hod formation and one at the base of the overlying maastrichtian tor formation boundary. although it is possible that the ‘intra-hod’ unconformity may correspond in part to the mid-campanian hiatus in jylland, this event in the danish central graben has typically been attributed a santonian age (e.g. abramovitz et al. 2010). it is noteworthy, however that a mid-campanian inversion event is well-known from the central graben in the netherlands sector (van der molen et al. 2005). the ‘base-tor’ unconformity of the danish central graben may correspond to the base maastrichtian unconformity recognised in this study. future work will focus on tying the jylland data set with that of the betterstudied chalk group of eastern denmark. acknowledgements the regional investigations to assess the geothermal energy potential in denmark are supported by the danish agency for science, technology and innovation, and interreg4a (european regional development fund). references abramovitz, t., andersen, c., jacobsen, f.c., kristensen, l. & sheldon, e. 2010: 3d seismic mapping and porosity variation of intra-chalk units in the southern danish north sea. in: vining, b.a. & pickering, s.c. (eds): petroleum geology: from mature basins to new frontiers 1, 537– 548. london: geological society. esmerode, e.v., lykke-andersen, h. & surlyk, f. 2007: ridge and valley systems in the upper cretaceous chalk of the danish basin: contourites in an epeiric sea. in: viana, a.r. & rebesco, m. (eds): economic and palaeoceanographic significance of contourite deposits. geological society special publication (london) 276, 265–282. gearhart 1986: biostratigraphy report. texaco 5410/5-1 danish well [kegnæs-1], 33 pp. unpublished report, gearhart geo consultants ltd. (in archives of geus – file 9689). liboriussen, j., ashton, p. & tygesen, t. 1987: the tectonic evolution of the fennoscandian border zone in denmark. tectonophysics 137, 21–29. lykke-andersen, h., & surlyk, f. 2004: the cretaceous–palaeogene boundary at stevns klint, denmark: inversion tectonics or sea-floor topography? journal of the geological society (london) 161, 343–352. rasmussen, s.l. & surlyk, f. 2012: facies and ichnology of an upper cretaceous chalk contourite drift complex, eastern denmark, and the validity of contourite facies models. journal of the geological society (london) 169, 435–447. surlyk, f. & lykke-andersen, h. 2007: contourite drifts, moats and channels in the upper cretaceous chalk of the danish basin. sedimentology 54, 405–422. thomsen, e. 1995: kalk og kridt i den danske undergrund. in: nielsen, o.b. (ed.): danmarks geologi fra kridt til i dag, 31–67. århus: geologisk institut, aarhus universitet. van der molen, a.s., dudok van heel, h.w. & wong, t.e. 2005: the influence of tectonic regime on chalk deposition: examples of the sedimentary development and 3d-seismic stratigraphy of the chalk group in the netherlands offshore. basin research 17, 63–81. authors’ addresses c.l. & l.o.b., department of geosciences and natural resource mangement, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: gnk171@alumni.ku.dk j.i., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. geological survey of denmark and greenland bulletin 41, 2018, 63-66 63 the geology of the paleoproterozoic karrat group in west greenland (71°–74°50´n) was investigated during the field seasons 2015–2017, using a combination of digital photogrammetry and traditional field work in a collaboration between the geological survey of denmark and greenland and the ministry of minerals resources of greenland. the area is characterised by steep alpine terrain with more than 2000 m of relief that in many places is completely inaccessible, which makes field work extremely difficult. therefore 3d mapping using digital photogrammetry is an invaluable tool in the investigation of the region. early geological investigations of the area involved the first use of photogrammetry in greenland (henderson & pulvertaft 1987). this contribution serves as an example of the present-day use of photogrammetry in geological interpretation, following the workflow outlined in sørensen & dueholm (2018). during the last three years, more than 50  000 stereo images have been collected using handheld, calibrated digital cameras while conducting field work in the area (rosa et al. 2016, 2017, 2018). the images, which cover large parts of the steep cliff sections in which the geology is superbly exposed, are essential to the ongoing revision of the geological map sheets covering the area. here we present a small subset of the data from the island of karrat (fig. 1), showcasing the potential of 3d geological mapping in greenland as well as presenting new insights into the geology of the karrat group. regional geology the karrat region is part of the rinkian fold belt of west greenland (henderson & pulvertaft 1967). the area mainly consists of reworked archaean gneisses overlain by supracrustal successions of the palaeoproterozoic karrat group. the group initially comprised two formations: the qeqertarssuaq and nûkavsak formations, but was later extended to also include the mârmorilik formation (henderson & pulvertaft 1987), originally considered to be of archaean age but later shown to be palaeoproterozoic and resting with a depositional unconformity on archaean gneiss (garde 1978). the karrat group and its archaean basement were metamorphosed and folded during the rinkian orogeny c. 1.9–1.8 ga (henderson & pulvertaft 1987; grocott & pulvertaft 1990). data acquisition and preparation stereo images were collected with calibrated, hand-held digital slr cameras from a boat (which served as base-camp) and from a helicopter used for day excursions and to supremote geological mapping using 3d photogrammetry: an example from karrat, west greenland erik vest sørensen and pierpaolo guarnieri 750 85 0 800800 900 750 700 750 8 0 0 9 00 750 850 850 800 800 750 700 150 750 50 100 50 100150 700 250 200 300 350 400 450 650 600550 500 28 73 7 44 62 11 14 28 47 15 51 60 45 14 28 31 60 21 52 50 52 86 13 14 50 43 56 53 38 61 38 40 45 41 44 44 48 46 58 39 42 greenland k kangilleq fm nûkavsak fm karrat group umanak gneiss mica schist qeqertarssuaq fm quartzite pegmatite foliation40 karrat 53°w 53°03'w53°06'w 71 °3 0' n 71 °2 9' n 1 km fig . 2 b fig. 2c fig. 1. part of the new geological map of karrat island (location in greenland marked on inset map), prepared as part of an ongoing revision of the 1:100 000 scale regional geological map sheets. the digital elevation model was generated from the oblique stereo-images collected during field work. inset boxes show the approximate positions of figs 2b, c. © 2018 geus. geological survey of denmark and greenland bulletin 41, 63–66. open access: www.geus.dk/bulletin http://www.geus.dk/bulletin 6464 port field camps. the images were typically collected while moving along the cliff faces in straight or gently curving trajectories tens to hundreds of kilometres long at varying distances to the cliffs. we here present results from a subset of the images collected from boat and helicopter flights around karrat. we used a hand-held nikon d800e (36 megapixel) digital, single lens reflex camera equipped with a carl zeiss distagon 35 mm lens that was pointed perpendicularly to the slope of interest through an open helicopter window or from boat. with the images we have almost complete coverage of the island with a resolution of c. 0.1–0.5 m (pixel size on the ground). the images were prepared for 3d mapping following the exact workflow of sørensen & dueholm (2018). gps positional data collected together with the images during field work were used as a first approximation for absolute positioning. the absolute orientation was subsequently refined through a proper bundle adjustment also including pass points measured stereoscopically in monochrome aerial photographs on a scale of 1:150 000 (i.e. points also identified in the oblique stereo-images), as well as planar levelling points (sea-level points) measured in the oblique images. the absolute accuracy of the stereoscopic models is around 3 m (xyz) while the photogrammetric or relative accuracy is in the millimetre to centimetre range. the images were subsequently used to extract elevation data for karrat, using sure software from nframes. a digital elevation model of karrat island with a 2 × 2 m grid (fig. 1) was produced to assist in the geological interpretation, as well as a so-called point cloud (a set of data points in 3d space representing the terrain surface) of the island to be used for visualisation purposes (fig. 2a). data interpretation – 3d mapping with the images properly oriented in 3d, essentially all corners of the island can be visited stereoscopically with the ease of a mouse-click and a geological interpretation can be performed. in this way it is possible, so to speak, to bring the rock exposures into the laboratory where the geology can be analysed. whereas previous geological investigations have taken place in the more accessible lower parts of karrat, here we focus on the inaccessible, higher parts of the island. karrat represents the north-westernmost exposure of the kigarsima nappe (henderson and pulvertaft 1987). the southwestern top of the island displays the overturned basement core of this tectonic nappe that forms an inverted sequence of archaean banded gneiss (umanak gneiss) with amphibolite layers sitting structurally on top of garnet-mica schist and quartzite of the qeqertarssuaq formation, metavolcanic rocks of the kangilleq formation and biotite schist of the nûkavsak formation (fig. 2). in addition to the actual 3d mapping with tracing of geological units, bedding and foliation (figs 2b, c), a powerful feature of the 3d mapping tool is the possibility to gather structural data remotely. this makes it possible to extend structural information from the shoreline, from where most structural data are usually collected during field work, up to outcrops at the top of the mountain (fig. 2a). this is important as it gives a more complete data coverage, and because the karrat region is structurally complex with multiple deformation stages (henderson & pulvertaft 1987; grocott & mccaffrey 2017). the structural data obtained with 3d photogrammetry consist of strike and dip of bedding/foliation, faults and thrusts together with traces of the geological boundaries between lithological or lithostratigraphic units. the data are stored as points (vertexes) along vectorised lines, so-called polylines. the strike and dip measurements presented here were calculated for each vertex of the polylines obtained through the 3d mapping as a moving average of best fitted planes by least square adjustment. more specifically, for each individual vertex a search window including seven adjacent vertexes was used in the calculation. in this way strike and dip measurements were generated for all vertexes of the mapped polylines. the measurements were subsequently filtered based on the standard deviation of each measurement. the result of this is a dataset of georeferenced points with calculated strikes and dips of their associated planar surfaces, which can be plotted on a geological map and analysed using stereoplots (fig. 2). the karrat island dataset consists of foliation in the umanak gneiss and amphibolites, bedding/ foliation in quartzites of the qeqertarssuaq formation, bedding of folded metavolcanic rocks of the kangilleq formation and the intensely folded metagreywacke strata of the nûkavsak formation, as well as folded pegmatites. fig. 2. point cloud model of karrat, looking north. a: lower hemisphere stereographic plots 1–4 of foliation and bedding measurements obtained from 3d polylines and shown as pole-to-bedding great circles and density averages. average fold axis orientations are defined by the intersection of great circles, shown with stars on the individual stereoplots. b: structural data from the southern part of the karrat showing the overturned limb of the kigarsima nappe (umanak gneiss and qeqertarssuaq formation), folded metavolcanic rocks of the kangilleq formation and nûkavsak formation metagreywackes. c: detailed close-up of the intense kink folds (purple lines) in the nûkavsak formation (numbers indicate dip direction/dip angle. blue line at the base: thrust fault dipping 30°nw compatible with the average fold trends. red line: normal fault dipping 60°w with a measured offset of 6 m. 65 d en sit y (% ) 0 6 12 18 24 30 10 0 m 30 0 m 93 4 m 78 4 m 78 5 m 91 2 m 77 6 m a b c w sw e n e 2b 2c ka ng ille q fm ka ng ille q fm n ûk av sa k fm th ru st (3 48 /3 0) fo ld a xi s (2 50 /1 0) 7 76 m 7 84 m 7 85 m 34 0/ 32 fau lt ( 26 4/6 0) o ffs et 6 m n ûk av sa k fm n ûk av sa k fm u m an ak g ne iss q eq er ta rs su aq f m (n = 1 00 )1 ka ng ille q fm (n = 3 5) 2 pe gm at ite s ( n = 35 ) 3 n ûk av sa k fm (n = 1 30 ) 4 av er ag e fo ld a xi s fo ld a xi s o f f ig. 2 c 6666 previously, grocott and mccaffrey (2017) described the emplacement of the kigarsima nappe towards ene as established by structural analysis of stretching lineations along a basal thrust contact. the authors also described an intersection lineation between cleavage and bedding, gently plunging towards sw, which appears to be compatible with a stretching lineation defined by hornblende and biotite minerals oriented wnw−ese that they related with a later top-to-nw tectonic transport overprinting the kigarsima structures. our new data presented in fig. 2 are in good agreement with the observations of grocott and mccaffrey (2017). at karrat island only the lower limb of the kigarsima nappe is preserved, represented by the umanak gneiss and qeqertarssuaq formation (fig. 2b). the c. 100 new foliation measurements (fig. 2 stereoplot 4) calculated from the mapped 3d polylines (fig. 2b) describe the geometry of a large recumbent fold with a subhorizontal, nnw−sseto nw−se-trending fold axis. this structural trend is compatible with ene−wsw to ne−sw compression that is consistent with the eneto ne-tectonic transport suggested by grocott and mccaffrey (2017). in contrast, a different structural trend is observed in the younger lithostratigraphic units. from the structural data (fig. 2a stereoplots 1–3) it appears that the structural trends are rotated almost 90°. in fact, the 40 measurements obtained from the folded metavolcanic rocks of the kangilleq formation (fig. 2a stereoplot 2) located in the overturned limb of the kigarsima nappe show an average ne-plunging fold axes that is similar to the fold axes obtained from the 130 measurements (fig 2a stereoplot 4) from the large kink folds observed in the nûkvsak formation (fig. 2c). these trends of folds are compatible with nw−se compression that in turn seems to be consistent with the top-to-nw tectonic transport indicated by grocott and mccaffrey (2017). similar fold axis trends are observed in the mapped pegmatites that cut the stratigraphy in the central part of the island (figs 1 and 2a stereoplot 3). this establishes an important cross-cutting relationship between the folding event and the pegmatites which are probably related to high temperature metamorphism dated at c. 1830 ma (rosa et al. 2017; kirkland et al. 2017). summary this study demonstrates how 3d mapping can be used for geological mapping in remote and inaccessible areas such as greenland following the procedures of sørensen & dueholm (2018). this is done with an example from the island of karrat, west greenland. using just a digital camera, we have generated a new revised geological map of karrat including topography, geology and structural data. our approach and methodology offer great support to standard field work where detailed outcrop information can be extended into regional-scale mapping. acknowledgments this work was carried out within the framework of an ongoing project financed by the geological survey of denmark and greenland and the ministry of mineral resources of greenland. we thank asger ken pedersen and ken mccaffrey for the helpful comments and suggestions. references garde, a.a. 1978: the lower proterozoic mârmorilik formation, east of mârmorilik, west greenland. meddelelser om grønland 200(3), 71 pp. grocott, j. & pulvertaft, t.c.r. 1990: the early proterozoic rinkian belt of central west greenland. in: lewry, j.f. & stauffer, m.r. (eds): the early proterozoic trans-hudson orogen of north america. geological association of canada, special paper 37, 443–463. grocott, j. & mccaffrey, k.j.w. 2017: basin evolution and destruction in an early proterozoic continental margin: the rinkian fold–thrust belt of central west greenland. journal of the geological society (london) 174, 453–467. henderson, g. & pulvertaft, t.c.r. 1967: the stratigraphy and structures of the precambrian rocks of the umanak area, west greenland. meddelelser dansk geologisk forening 17, 1–20. henderson, g. & pulvertaft, t.c.r. 1987: geological map of greenland, 1:100 000, mârmorilik 71 v.2 syd, nûgâtsiaq 71 v.2 nord, pangnertôq 72 v.2 syd. descriptive text. 72 pp., 8 plates. copenhagen: geological survey of greenland. kirkland, c.l., hollis, j., danišík, m., petersen, j., evans, n. j., & mcdonald, b.j. 2017: apatite and titanite from the karrat group, greenland; implications for charting the thermal evolution of crust from the u-pb geochronology of common pb bearing phases. precambrian research 300, 107–120. rosa, d., guarnieri, p., hollis, j., kolb, j., partin, c., petersen, j., sørensen, e.v., thomassen, b., thomsen l. & thrane, k. 2016: architecture and mineral potential of the paleoproterozoic karrat group, west greenland. danmarks og grønlands geologiske undersøgelse rapport 2016/12, 98 pp. rosa, d., dewolfe, m., guarnieri, p., kolb, j., laflamme, c., partin, c., salehi, s., sørensen, e.v., thaarup, s., thrane, k. & zimmermann, r. 2017: architecture and mineral potential of the paleoproterozoic karrat group, west greenland – results of the 2016 season. danmarks og grønlands geologiske undersøgelse rapport 2017/5, 98 pp. rosa, d., bernstein, s., dewolfe, m., dziggel, a., grocott, j., guarnieri, p., kolb, j., partin, c., sørensen, e.v. & zimmermann, r. 2018: architecture and mineral potential of the paleoproterozoic karrat group, west greenland – results of the 2017 season. danmarks og grønlands geologiske undersøgelse rapport 2018/23, 102 pp. sørensen, e.v. & dueholm, m: 2018: analytical procedures for 3d mapping at the photogeological laboratory of the geological survey of denmark and greenland. geological survey of denmark and greenland bulletin 41, 99–104 (this volume). authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: evs@geus.dk. mailto:evs@geus.dk geological survey of denmark and greenland bulletin 31, 2014, 15-18 15 borehole logs from the precambrian basement on bornholm, eastern denmark: geology and groundwater flow peter gravesen, bertel nilsson, per rasmussen and stig a. schack pedersen bornholm is situated in the sorgenfrei–tornquist zone that separates the north-west european craton from the baltic shield and the east european platform (fig.1). the precambrian basement of northern and eastern bornholm consists of different granitic and gneissic mesoproterozoic rocks that are dated to c. 1455 ma (waight et al. 2012). it appears from the age data that granitic magmatism, deformation and metamorphism occurred over a relatively short time period. the rocks contain abundant pegmatite and aplite bodies. more than 250 mafic dykes occur. the dykes were intruded during three precambrian events at c. 1326 ma, 1220 ma, 950 ma and during a permian event at c. 300 ma (holm et al. 2010). the present study focuses on the østermarie–paradisbakke area north of paradisbakkerne and deals with the paradisbakke migmatite and part of the bornholm gneiss. the aim of the study was to map the distribution of fractures in the rocks and determine the groundwater flow in these low-permeability rocks using outcrop data and borehole logs. the survey was part of an investigation of potential areas for disposal of radioactive waste from the research centre risø area (gravesen et al. 2011a, b, 2012, 2013). rocks, fractures and groundwater the two types of crystalline basement rocks in the study area show similar mineralogy and are dominated by k-feldspar (35–38%), quartz (23–30%), plagioclase (22–25%) and hornblende (11–17%); biotite and other minerals also occur (micheelsen 1961). the bornholm gneiss is medium-grained and grey or reddish grey. the fineto medium-grained paradisbakke migmatite consists of almost parallel, light grey © 2014 geus. geological survey of denmark and greenland bulletin 31, 15–18. open access: www.geus.dk/publications/bull skåne bornholm bal t i c sea faults paradisbakkerne østermarie a b cretaceous clay, sand and limestone jurassic sand and clay silurian shale cambrian sandstone granite and gneiss with dykes 25 km n balt ic sh ie ld east european plat form nw european craton s–t zone bornholm fig. 1. geological map of bornholm and skåne in southern scandinavia. the location of the two cross-sections shown in fig. 4 is indicated. inset: bornholm’s location in the sorgenfri–tornquist zone (s–t zone; from graversen 2009). fig. 2. paradisbakke migmatite in præstebo quarry at the northern rim of paradisbakkerne. horizontal and vertical fractures are common whereas oblique fractures are rare. the height of the outcrop is c. 12 m. 1616 granitic quartz-feldspar veins in a darker matrix. both rocks contain pegmatite bodies and are cut by thin nne–ssworiented dykes that were intruded c. 1220 ma ago. minor bodies of reddish grey granite are also present. the rocks are cut by four fracture systems, which can be observed in quarries with paradisbakke migmatite along the northern rim of paradisbakkerne (von bubnoff 1942; fig. 2). after deformation and cooling in the precambrian, large tectonic faults and fractures and minor vertical fractures were formed. the main fracture system in the area has an nne–ssw orientation, corresponding to the direction of the mafic dykes (micheelsen 1961). large linear, fault-controlled valleys are found in paradisbakkerne and in the western part of the precambrian basement on bornholm, but such valleys are absent in the study area. however, small vertical fractures with the same orientation are seen in outcrops. a second ese–wnw-oriented verical tectonic fracture system that shows the same orientation as palaeozoic faults was probably formed during a main phase of wrench faulting in the sorgenfrei–tornquist zone in the late palaeozoic or later (graversen 2009). zones of 4.5 m thickness contain thin fractures with centimetre-sized spacing, while single fractures can have up to 5 m spacing. small faults with slickensides are present and show that some horizontal movement has occurred. in addition to the vertical or subvertical fractures, rare occurrences of oblique fractures and conjugate fracture sets are also found. a third system of horizontal fractures (sheet jointing) at levels from 25 m to 100 m depth was probably formed by load release when an overburden of supposed younger deposits, probably of cambrian to quaternary age, was eroded and removed (spencer 1969). we assume that the spacing 10 20 30 40 50 0 100 10000 10050 inf low zo ne lit ho log y dep th (m ) flo w lo g (% ) major minor major major major precambrian gneiss and migmatitequaternary clayey till rela tiv e i nfl ow 400 natu ral ga mma (a pi) 800 fo rm ati on re sis tiv ity (o hm -m ) so nic ve loc ity (k m/s) 8040 con du cti vit y (m s/m ) 84 fig. 3. lithology and logs in borehole dgu no. 247.458. the flow-log data were acquired at a pumping rate of 1.3 m3 h–1. api: american petroleum institute units. 17 between the fractures increases downwards, because this is seen in quarries. a fourth fracture system consists of horizontal fractures near the ground surface with a spacing of 1–2 m, which also increases downwards. crushed rocks with blocky structures also occur. this system may have been formed by glaciers during the quaternary. the lower limit of the fractures is unknown but some horizontal fractures are found at 90 m below the ground surface. this means that water flow in vertical fractures that cross-cut the deep horizontal fractures can supply the deep fracture system with groundwater from above. the rocks in the area are slightly weathered and fe-containing minerals are oxidised to yellowbrown clayey, fe-rich deposits on some fracture surfaces. such deposits may prevent water flow in some of the fractures. most of the area is covered by a clayey till up to 6 m thick. the groundwater reservoir is characterised by a network of vertical to subvertical and horizontal to subhorizontal fractures. surface water is mainly transported to the groundwater zone in the vertical fractures and the groundwater can be transported over long distances in the horizontal fractures. it is difficult to map the subsurface fractures but data from borehole logs can contribute to our understanding of the network. normally the groundwater table is found a few metres below the ground surface but pumping will often lower the groundwater table because the storage capacity in the reservoir is small. the yield of boreholes in basement rocks on bornholm is commonly low, but boreholes supplying østermarie waterworks show high yields because of fracture systems in the rocks in this area. borehole logs eight private water boreholes in the study area were investigated by geophysical wireline logging. the water pump was removed in the morning and re-installed in the late afternoon after logging. the following geophysical parametres were measured: natural gamma radiation, formation resistivity, sonic velocity, conductivity and impeller flow. conductivity and flow log data were acquired during groundwater pumping from the borehole. natural gamma, resistivity and sonic velocity logs provide information on lithological variation in the borehole. the conductivity log gives information about groundwater chemistry and under certain conditions also provides information on groundwater inflow zones. the flow log measures the vertical flow velocity in the borehole. changes in flow velocity during pumping indicate groundwater inflow zones at specific depths or depth intervals in the borehole. results the results from the logging of borehole dgu no. 247.458 are shown in fig. 3 and two constructed geological sections through the study area are shown in fig. 4. in the c. 5 km long north–south section, the fracture inflow into four boreholes is shown (fig. 4a). the depths of water flow in the basement rocks indicate the location of fractures and the inflow is evaluated semi-quantitatively as ‘major’ or ‘minor’ in each borehole. ‘major’ indicates that the fracture inflow contributes a large part of the total inflow whereas ‘minor’ indicates only a minor contribution. ‘major’ inflows are interpreted from the flow log where an increase in the flow log (%) occurs, measured from the bottom towards the top of the borehole. ‘minor’ inflows are interpreted from the conductivity log where –25 –25 0 25 50 75 el ev at io n (m a .s. l.) 1 km south north groundwater level sea 247.369 247.403 247.567 247.541 247.432 a groundwater level 247.496 247.403 247.458 247.661 0 25 50 100 west east 1 kmb major inflow minor inflow fig. 4. two cross sections through the logged boreholes. for location see fig. 1. note different horizontal and vertical scales. 1818 changes in conductivity appear. the conductivity log can be more sensitive to minor influx zones than the flow log. it is relevant to compare the boreholes because they have the same low specific yield (0.05–0.22 m3/h–1). figure 4 shows fracture inflow at four different depth intervals. major inflow is registered near the bottom of the boreholes and minor fracture inflow is seen closer to the surface. the groundwater table is situated a few metres below the surface in all boreholes. it is not possible to directly determine the orientation of the fractures, but fractures that are found at the same level, at c. 10 to 12 m a.s.l., are likely to be horizontal or subhorizontal and hydraulically connected. in the c. 8 km long west–east section, fracture inflow is registered between c. 100 m a.s.l. and 25 m b.s.l. (fig. 4b). borehole dgu no. 247.496 at østermarie waterworks is included in this section (rasmussen et al. 2007). inflow is seen in all boreholes but the largest inflow is found at the bottom of borehole dgu no. 247.403. several important waterbearing fracture zones are found in the waterworks borehole. the fractures occur at shallower depths at at least four levels. the results from the logging survey appear to confirm the occurrence of important fracture systems at depth. fractures are located at different levels. horizontal fracture systems seen over a large area are found at 10–12 m a.s.l. and a similar system occurs at c. 20 m b.s.l. several smaller fracture systems without large water flow also occur. the most important and largest fracture system is found at 90 m below the surface, usually at the bottom of the boreholes. as mentioned above we cannot determine the orientation of the fractures from the borehole log survey, but the structural model indicates major elements of the fracture systems. the horizontal fracture spacing apparently increases with depth, but it is still posssible to find fracture systems with major water flow if they are connected to other vertical and horizontal fractures. conclusions the conclusions of the investigations are as follows: 1. the logging of the boreholes shows that fracture systems occur in all of them. 2. the fractures occur down to at least 90 m below the surface, and flow logging documents the occurrence of groundwater flow down to this depth. the downward transport of surface water to the horizontal fractures may be via vertical fractures to at least the same depth. 3. horizontal fractures are found at 10–12 m a.s.l. and at 20 m b.s.l. the fractures appear to be connected over kilometre-long distances. acknowledgements the danish parliament is thanked for financial support. we are grateful to local landowners that gave us permission to log their private boreholes. references graversen, o. 2009: structural analysis of superposed fault systems of the bornholm horst block, tornquist zone, denmark. bulletin of the geological society of denmark 57, 25–49. gravesen, p., nilsson, b., pedersen, s.a.s. & binderup, m. 2011a: low and intermediate level radioactive waste from risø, denmark. location studies for potential disposal areas. report no. 4. characterisation and description of areas bornholm. danmarks og grønlands geologiske undersøgelse rapport 2011/44, 85 pp. gravesen, p., binderup, m., nilsson, b. & pedersen, s.a.s. 2011b: geological characterisation of potential disposal areas for radioactive waste from risø, denmark. geological survey of denmark and greenland bulletin 23, 21–24. gravesen, p., nilsson, b., binderup, m. larsen, t. & pedersen, s.a.s. 2012: lavog mellem radioaktivt affald fra risø, danmark. omegnsstudier. rapport nr. 1. område østermarie-paradisbakkerne, bornholms regionskommune. danmarks og grønlands geologiske undersøgelse rapport 2012/123, 100 pp. gravesen, p., nilsson, b., binderup, m., larsen, t.b. & pedersen, s.a.s. 2013: geology, seismic activity and groundwater conditions at six potential disposal sites for radioactive waste from risø, denmark. geological survey of danmark and greenland bulletin 28, 13–16. holm, p.m., pedersen, l.e. & højsteen, b. 2010: geochemistry and petrology of mafic proterozoic and permian dykes on bornholm, denmark: four episodes of magmatism on the margin of the baltic shield. bulletin of the geological society of denmark 58, 35–65. micheelsen, h. 1961: bornholms grundfjæld. meddelelser fra dansk geologisk forening 14, 308–349. rasmussen, p., klitten, k., nielsen, s. & jensen, p. 2007: bornholms regionskommune. logging og vandkemi i vandforsyningsboringer, 2006. danmarks og grønlands geologiske undersøgelse rapport 2007/36, 91 pp. spencer, e.w. 1969: introduction to the structure of the earth, 597 pp. new york: mcgraw-hill. von bubnoff, s. 1942: beiträge zur tektonik des skandinavischen südrandes. 2. die älteren granite bornholms im rahmen der svekofennidischen tektogenese. neues jahrbuch für mineralogie, geologie und paläontologie, beilagen-band 87, 277–396. waight, t., frei, d. & storey, m. 2012: geochronological constraints on granitic magmatism, deformation, cooling and uplift on bornholm, denmark. bulletin of the geological society of denmark 60, 23–46. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: pg@geus.dk mailto:es@geus.dk geological survey of denmark and greenland bulletin 28, 2013, 1-8 1 geological survey of denmark and greenland bulletin 28 • 2013 review of survey activities 2012 edited by ole bennike, adam a. garde and w. stuart watt geological survey of denmark and greenland ministry of climate, energy and building 22 geological survey of denmark and greenland bulletin 28 keywords geological survey of denmark and greenland, survey organisations, current research, denmark, greenland. cover photographs from left to right 1. the 1097 m high finnefjeld is located in the central part of the maniitsoq impact structure, west greenland. photograph: adam a. garde. 2. investigation of a coastal cliff section at a potential disposal site for radioactive waste. photograph: merete binderup. 3. setting up an automatic camera to monitor frontal variations of a glacier in west greenland. photograph: robert fausto. 4. collecting samples in cameroon, using a hand auger. photograph: christian knudsen. frontispiece: facing page in 2012 geus published a geological map of the island of mors in north-west jylland. the map sheet comprises the coastal cliff section hanklit on northern mors, with exposures of light layers of moler and dark layers of volcanic ash, overlain by glaciofluvial deposits from the last ice age. the moler has been folded and up-thrusted by ice advancing from the north. photograph: stig a. schack pedersen. chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, department of geoscience, aarhus university; minik rosing, geological museum, university of copenhagen; finn surlyk, department of geosciences and natural resource management, university of copenhagen scientific editors: ole bennike, adam a. garde and w. stuart watt editorial secretaries: jane holst and esben w. glendal referees (numbers refer to first page of reviewed article): holger lykke andersen (21); anonymous (37); luke b. bateson, uk (25); stefan bernstein, dk (45); albertas bitinas, lt (21); kristian bitsch, dk (33); dave burgess, ca (69); jakob qvortrup christensen, dk (13); gregers dam, dk (61); ole gravesen, dk (65); christoph a. hauzenberger, g (53); jens havskov, no (41); jens asger jensen, dk (33); maths halstensen, no (37); larry hulbert, ca (45); joakim stiel korshøj, dk (25); john korstgård, dk (65, 73); peter langen, dk (69); nicolaj krog larsen, dk (29); poul-henrik larsen, dk (61); björn lund, se (41); mogens marker, no (57); claudio milisenda, g (53); matthias moros, g (17); ole bjørslev nielsen, dk (9); asger ken pedersen, dk (49); anette petersen, dk (29); alar rosentau, ee (17); martin sønderholm, dk (9); inga sørensen, dk (13); henrik stendal, gl (49, 73); jeroen van gool, dk (57); thomas zack, se (73). illustrations: benny m. schark, jette halskov, stefan sølberg, kristian a. rasmussen, willy l. weng, frants v. platen-hallermund layout and graphic production: kristian a. rasmussen printers: rosendahls-schultz grafisk a/s, albertslund, denmark manuscripts received: 22 january 2013 – 14 march 2013 final versions approved: 15 february 2013 – 21 may 2013 printed: 10 july 2013 issn 1604-8156 isbn 978-87-7871-357-5 citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 28, 76 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk and at www.geus.dk/publications/bull © de nationale geologiske undersøgelser for danmark og grønland (geus), 2013 for the full text of the geus copyright clause, please refer to www.geus.dk/publications/bull 3 44 tanzania mozambique zambia indonesia vietnam china korea greenland philippines brazil denmark ghana nigeria cameroon 7 review of survey activities 2012 f.g. christiansen 9 drowning of the miocene billund delta, jylland: land–sea fluctuations during a global warming event e.s. rasmussen, t. utescher and k. dybkjær 13 geology, seismic activity and groundwater conditions at six potential disposal sites for radioactive waste from risø, denmark p. gravesen, b. nilsson, m. binderup, t.b. larsen and s.a.s. pedersen 17 a baltic ice lake lowstand of latest allerød age in the arkona basin, southern baltic sea o. bennike and j.b. jensen 21 late glacial to early holocene development of southern kattegat c. bendixen, j.b. jensen, o. bennike and l.o. boldreel 25 terrain subsidence detected by satellite radar scanning of the copenhagen area, denmark, and its relation to the tectonic framework p. r. jakobsen, u. wegmuller, r. capes and s.a.s. pedersen 29 geological map of denmark 1:50 000 – map sheet mors, nw denmark s.a.s. pedersen, p.r. jakobsen, l.tougaard and p. gravesen 33 assessing urban groundwater table response to climate change and increased stormwater infiltration m.t. randall, l. troldborg, j.c. refsgaard and j.b. kidmose 37 evaluation of total groundwater abstraction from public waterworks in denmark using principal component analysis b.l. sørensen and r.r. møller 41 seismic activity in denmark: detection level and recent felt earthquakes t. dahl-jensen, p.h. voss, t.b. larsen and s. gregersen 5 tanzania mozambique zambia indonesia vietnam china korea greenland philippines brazil denmark ghana nigeria cameroon geus working areas 2012. orange areas are covered in this volume. 45 the norite belt in the mesoarchaean maniitsoq structure, southern west greenland: conduit-type ni-cu mineralisation in impact-triggered, mantlederived intrusions? a.a. garde, j. pattison, t.f. kokfelt, i. mcdonald and k. secher 49 geochemistry and petrology of gold-bearing hydrothermal alteration zones on qilanngaarsuit, southern west greenland m. koppelberg, a. dziggel, d.m. schlatter, j. kolb and f.m. meyer 53 fingerprinting of corundum (ruby) from fiskenæsset, west greenland n. keulen and p. kalvig 57 lineament mapping and geological history of the kangerlussuaq region, southern west greenland k.e.s. klint, j. engström, a. parmenter, t. ruskeeniemi, l.c. liljedahl and a. lehtinen 61 calibration of spectral gamma-ray logs to deltaic sedimentary facies from the cretaceous atane formation, nuussuaq basin, west greenland g.k. pedersen, n.h. schovsbo and h. nøhr-hansen 65 a new seamless digital 1:500 000 scale geological map of greenland m. pedersen, w.l. weng, n. keulen and t.f. kokfelt 69 darkening of the greenland ice sheet due to the melt-albedo feedback observed at promice weather stations d. van as, r.s. fausto, w.t. colgan, j.e. box and the promice project team 73 titanium minerals in cameroon c. knudsen, j. penaye, m. mehlsen, r.k. mclimans and f. kalsbeek 66 7 review of survey activities 2012 flemming g. christiansen deputy director 2012 was a good and stable year for the geological survey of denmark and greenland (geus). in recent years geus has been through a long – and very constructive – process of establishing a new strategy that reflects the changes in society and new demands from many different stakeholders. with a new strategy in place there has been greater focus on geus’ activities and research projects. 2012 was a very active year with many projects, field work and offshore data acquisition, which promises well for maintaining a high level of research in the coming years. with the establishment of the new series geological survey of denmark and greenland bulletin in 2003 it was decided to make a yearly review of survey activities. this issue is the tenth and, together with previous issues, provides a good overview of the survey’s wide range of research and advisory activities. this issue contains a total of 17 four-page papers, nine on denmark, seven on greenland and one on a project in cameroon. activities in denmark the activities and research in denmark by geus cover many different topics within our main programme areas: data, water, energy, mineral resources as well as nature and climate. as a follow-up on many previous studies of the miocene succession and its groundwater resources, one paper gives a detailed discussion of the drowning of the billund delta in jylland during a period of previous global warming. geus has been involved in the technical work that is required before a permanent disposal site can be selected for lowand intermediate-level radioactive waste. one paper describes the geological data and knowledge and key parameters such as lithology, groundwater, seismic activity, effect of climate change and local infrastructure that have been applied to narrow down the initial number of 22 areas to six potential sites. geus is involved in many studies of quaternary and recent geological processes. one paper describes deposits from the baltic ice lake in the arkona basin in the southern baltic sea. another paper presents results on the late glacial to early holocene development of southern kattegat. a third paper describes terrain subsidence interpreted from satellite radar scanning with special focus on copenhagen and its tectonic framework. systematic geological mapping on a scale of 1:25 000 and publication of 1:50  000 scale systematic sheets continues. in some places it is appropriate to make regional maps of, for example islands, on a single sheet. an example of this is described in a paper on the map sheet mors. geologically, mors is an interesting island with beautifully exposed glaciotectonic structures in cliff sections, and the characteristic eocene clayey diatomite with volcanic ash layers, a unit known as ‘moler’. the use of groundwater is very important for danish society and geus carries out many studies on water resources and possible future changes due to climate and use. one paper discusses models for urban groundwater table response to climate change and increased stormwater infiltration using the town of silkeborg and a proposed course of a new motorway around silkeborg as a case study. another paper evaluates groundwater abstraction from public waterworks in denmark by using advanced statistical methods to correct data. geus records seismological events at six locations in denmark. one paper describes the developments in detection level and completeness from 2000 to 2012 with examples of recent felt earthquakes (north sea on 19 february 2010 and kattegat on 6 august 2012) and explosions. activities in greenland there was a high level of field activities in greenland in 2012 with a major mapping and geochemical programme in south-east greenland, a large field and shallow-core drilling programme in north greenland, and the lomrog iii cruise in the arctic ocean. the latter was the last data acquisition of the danish continental shelf project. many other field studies were also carried out. results from these large and small projects will be presented in the coming years. in this issue results are presented from other completed and ongoing projects. one paper discusses the possibility that nickel and copper occurrences in the norite belt in the maniitsoq area could be impact-induced as they are found in a recently described giant and deeply eroded, very old (3 ga) impact structure © 2013 geus. geological survey of denmark and greenland bulletin 28, 7–8. open access: www.geus.dk/publications/bull 88 in this area. gold occurrences have been reported during geus expeditions to west greenland some years ago, and another paper provides more constraints on ore formation and fluid-rock interaction. corundum in gem quality (ruby) has been known since the 1960s near fiskenæsset in southern west greenland and it is possible that mining will start within a few years. one paper gives a detailed description of the geochemistry that can be used to fingerprint the greenlandic rubies. since 2008, the greenland analogue project (gap) has carried out a wide range of studies in an area near kangerlussuaq in west greenland to understand the many different processes that might take place if a deep geological repository for spent nuclear fuel is affected by glaciation. one paper focuses on structural investigations in this area, especially the mapping of lineaments and other late features. as a followup on several decades of studies in the nuussuaq basin in west greenland, a spectral gamma-ray characterisation has been made on the cretaceous atane formation to set up a better model for interpretation of lithology and grain-size variation in wells without cores. systematic geological overview mapping of greenland has been the backbone of the survey’s work in greenland over the past 65 years ago. in the period 1971–2004, 14 maps were published on a scale of 1:500 000. with the development of the internet, geological maps worldwide are rapidly changing from traditional paper sheets to digital publications. geus has produced a new seamless digital 1:500 000 scale map covering the whole of greenland. one paper describes the background, the digitisation and the geological harmonization involved. it is expected that such maps will be widely used in the future. the new map is an important geus contribution to the global onegeology portal. the important monitoring programme of the greenland ice sheet (promice) continuously supplies new crucial data and one paper discusses the darkening of the greenland ice sheet and the increase in surface melting. international activities internationally geus works in many different countries with a variety of projects. the last paper in this issue is about titanium minerals in cameroon and it focuses on understanding the primary formation, weathering and re-deposition of rutile in the area. geological survey of denmark and greenland bulletin 38, 2017, 73-76 73 where there scientific or political reasons behind denmark’s decision to establish its first seismological stations for earthquake monitoring? in a nation where earthquakes are few and of small magnitude, it is remarkable that since 1927 the geological survey of denmark and greenland (geus), and its predecessors back to the danish geodetic institute have recorded seismological events from permanent stations in denmark and greenland and shared data through international data centres. as early as 1907, on private initiative by e.g. harboe, a seismological station was established in godhavn, west greenland (harboe 1911). however, with time it became clear that the phase readings from the instruments lacked precision (fig. 1), and in 1912 monitoring was discontinued due to lack of funding (geodætisk institut 1978). having never officially joined the international seismological association, or suffered destructive earthquake within their territory, danish authorities did not at the time have any vested interest in operating expensive seismological stations. the origins of denmark’s seismological monitoring programme historians of science have argued that scientific knowledge of greenland equalled sovereignty (ries 2012; doel et al. 2016). inspired by their approach, i suggest that theformation of a network of danish and greenlandic seismological stations could be understood as an expression of national geopolitical strategy rather than a display of scientific vigour. by examining historical records from public institutions like the ministry of education, ministry of foreign affairs, ministry of war and the danish geodetic institute as well as letters from erik nørlund and inge lehmann, kept at the danish national archives and geus, it is possible to assess to what degree geopolitical considerations were a factor in establishing a seismological monitoring programme in denmark. in 1923, denmark became a full member of the international union of geodesy and geophysics (iugg), but it was not until october 1924 after the second iugg general assembly in madrid that (lack of) seismological monitoring from its territory became an issue for denmark. before the assembly, a national committee for seismology had hastily been formed under the leadership of niels erik nørlund, professor of mathematics at copenhagen university and director of the danish geodetic service (den danske gradmaaling). after the assembly, nørlund reported to the ministry of education that there had been considerable pressure on denmark to record and share geophysical data from greenland. so substantial was the pressure that nørlund warned that unless denmark initiated a programme of general geodetic research, in particular seismological monitoring and international data sharing, other nations would question denmark’s authority in greenland (danish national archives (dna) 1). it was a potent threat, since the danish government rightfully feared that their claim to sovereignty over greenland could be contested by their neighbours. for example, a stipulation of the 1917-treaty between denmark and the usa regarding the sale of the west indies to the usa was, in addition to paying dkk 25 million, that the usa also acknowledged denmark’s sovereignty over greenland. norway on the other hand made counterclaims to parts of eastern greenland in 1924 and 1931 (kragh et al. 2008). over the next months, nørlund worked on securing support for his envisioned network of seismic stations. as arctic geopolitics and the beginning of earthquake monitoring in denmark and greenland a. lif lund jacobsen fig. 1. seismograms from the old station in godhavn kept at geus. these seismograms from 1907–1912 are among the oldest seismological records in existence. photo: danish national archives. © 2017 geus. geological survey of denmark and greenland bulletin 38, 73–76. open access: www.geus.dk/publications/bull 7474 director of the danish geodetic service, nørlund was appointed to reorganise national geodesy, including triangulation efforts which traditionally had been a key interest area of the army’s department of topography (generalstabens topografiske afdeling). using his position, he added seismology to the list of intended activities. in order to secure the necessary political backing and financial support from private funds, nørlund employed several different arguments. in addition to demonstrating sovereignty over greenland, he argued that denmark needed seismological stations because all cultural and refined nations in europe had a least one. also scientifically, seismological data constituted a valuable source of information for geodetic research. finally, he argued that international data sharing was a means to maintain international peace and collaboration. during the winter 1924/25, enough political and financial support was mustered that nørlund and the danish geodetic service could begin building the network. establishing seismological stations from 1925 to 1927 from early on, it was clear that while the danish government was willing to authorise the construction of seismological stations, much of the funds to build stations and buy scientific instruments had to come from private donations. in the spring of 1925, orders for a variety of different seismographs and other instruments were placed at international suppliers in england, germany, russia and the usa. not knowing which seismic frequencies they could expect and at a time where no standards for performance and accuracy exsisted, orders were for state-of-the-art seismograph models of different mechanical or electrical designs. the costly instruments were all paid for by the carlsberg foundation which also made significant contributions towards the construction costs of the stations. in addition, carlsberg also agreed to provide an annual grant to cover operating costs for two of the three stations, a practice that continued until the 1950s. without any practical knowledge about the working of seismological instruments or how they would react to the extreme climate of greenland, it was decided that a station in denmark should act as a testing and comparison site for the different instruments. at the same time negotiations began to secure space for stations both in copenhagen and greenland. historical records show that geopolitical and strategic considerations played a significant role in where to locate the three planned stations (e.g. dna 2). being relatively close to the city, the permanent station in copenhagen (cop) was the first to be completed. in the spring of 1925, the army handed over two caponieres located in the old fortifications of copenhagen (vestvolden) to be used as a seismological station. still being part of the military defense system, access to the station was restricted until well after the second world war. installations began in the autumn of 1926. by 17 february 1927 all instruments were in place and monitoring began, but due to occasional failure of the timing, the working of the station was not considered satisfactory before march. hereafter publication of the records began in a seismological bulletin. by then cop was equipped with the following seismographs: a wiechert 1000 kg horizontal (fig. 2), a wiechert 1300 kg vertical, three galitzin instruments (two horizontal and a vertical), a 2-component milne-shaw. later a 2-component wood-anderson seismograph was added. today historical bulletins and log book from all the stations can be found at http://seis.geus.net/seismic_service. html. as the largest of the three seismological stations in the network, cop had a fulltime caretaker to do daily maintenance and recording, paid by the carlsberg foundation and living in a nearby cottage. establishing and installing denmark’s first permanent seismological station was not an easy task for the untrained personnel hired by nørlund, and in recognition of their valuable contribution, inge lehmann and two students received a bonus worth nearly two months’ salary (dna 3). by all accounts, this was also the later so famous seismologist inge lehmann’s first encounter with seismological work, which principles she first began to study in earnest later that year. the station in ivittuut (ivi), south west greenland, was the second to be completed (fig. 3). with the help fig. 2. in cop’s caponiére xiii, the purely mechanical wiechert horizontal seismograph is still in its original location. photo: casper brogaard højer. 75 of the danish mining company kryolith mine og handelsselskabet, a seismological vault was blasted into the bedrock about 250 m from the mine’s main blasting sites (!) to determine the effect of blasting on the seismological instruments tests were made with a portable seismograph during the summer of 1926, but it was not before the summer of 1927 that the intended seismographs, a wiechert vertical and a wiechert horizontal seismograph, were installed (dna 4). operated, for an extra fee, by the local radio-telegraphist employed by the mining company, the station began recording on 24 august 1927 (bulletin ivigtut 1929). when the third and final station in the danish network was constructed, it became clear that that geopolitical arguments worked both ways between arctic nations. in september 1925, nørlund wrote to the danish ministry of foreign affairs that the danish geodetic service was considering a permanent seismological station on jan mayen because the island had an active volcano. it was a delicate subject since norway’s meteorological institute had annexed part of the island on behalf of norway. denmark had refused to officially state its position on the matter but maintained that some buildings within the norwegian-claimed area were owned by denmark. with the aid of the ministry of foreign affairs, an agreement was made with norway in march 1926 to the effect that the danish geodetic service could place a station on the norwegian part of jan mayen. however, shortly afterwards in july, norway announced that their meteorological institute would expand its activities to the whole island and claim it on behalf of norway (dna 5). in the summer of 1926, initial seismological tests were carried out on jan mayen by the danish geodetic service, but results were poor and it was decided to establish the station at the newly founded colony of scoresbysund (ittoqqortoormiit). founded in 1925, one of the colony’s purposes was to establish a danish presence in east greenland where norway also had made claims. construction of the station (sco) began in the summer of 1927, also here funded by the carlsberg foundation. by then the danish geodetic service had enough experience with construction and installation to proceed relatively quickly. a cellar was blasted into the bedrock, and on top of that a low building was constructed. to protect the instruments from variations in temperature the cellar was covered by 80 tons of rock and access restricted to a low corridor one could only crawl through. a second building housed the radio station and electronic recording instruments (geodætisk institut 1930). the installation of the seismological instruments went smoothly, as the station was equipped with two galitzin horizontal seismographs and a galitzin vertical seismograph transferred from cop (den danske gradmaaling 1928). fully funded by the carlsberg foundation the station began operation on 12 january 1928, with the local radio-telegraphists in charge of daily recordings (fig. 4). in the beginning the paper seismograms were sent annually to the geodetic institute by ship, but later it became practice that results from large events were radioed immediately to copenhagen. after 1928 in 1928 the danish geodetic service and the department of topography of the army general staff (generalstabens topografiske afdeling) merged to become the danish geodetic institute (danmarks geodætiske institut). inge lehmann, who had been the principal figure in setting up the copenhagen station and analysing data from all the fig. 3. the entrance to the seismological vault in ivittuut is next to the old tennis court. today (2014), the court is unkempt and ivittuut largely deserted. fig. 4. inge lehmann (second from the left) inspects sco presumably in the summer of 1928. photo: from inge lehmann’s private archive at the danish national archives. 7676 stations, was appointed chief of the geodetic institute’s new seismic section. the value of the greenlandic seismological stations as evidence of sovereignty was soon tested. in 1931, norway claimed parts of east greenland, denmark opposed the claim and the matter was put to the permanent court of international justice in the hague. in preparation for the trial, the ministry of foreign affairs asked the danish geodetic institute to prepare a series of reports about their scientific activities in greenland. in july 1932, inge lehmann forwarded a special report about the seismological station in scoresbysund to the ministry (dna 5) which was submitted as part of the evidence supporting the danish claim. in april 1933, the court decided against norway, recognising danish sovereignty over all parts of greenland. scientifically, the three stations also soon proved valuable by providing quality seismic records from remote, low-noise greenland, with efforts put into timing records and adjustment of instruments in the challenging environment. it was partly data from the seismological stations in ivittuut and ittoqqortoormiit that in 1936 enabled inge lehmann to deduce the existence of the earth’s inner core (lehmann 1936). today geus records seismological data from about 25 locations in denmark and greenland, including the original stations in copenhagen, ivittuut and ittoqqortoormiit, now equipped with modern digital instruments. conclusions historical documents from the danish national archives and geus clearly show that international geopolitics and strategic considerations played a significant role in the location of the seismological stations. especially denmark’s need to express its sovereignty over greenland played a pivotal role. it was by invoking arguments of power, culture, science and international peace the director of the danish geodetic service, erik nørlund, was able to gain the necessary political and financial support for establishing a network of seismological stations, and in 1928 establish a permanent danish seismological monitoring authority under the danish geodetic institute. as the danish government’s main interest was the strategic value of the seismological stations, it fell to private donations to unlock the scientific potential of the stations. it was therefore the logistic support of the kryolith mine og handelsselskabet and the long-term financial commitment of the carlsberg foundation that made it possible to equip and operate the stations at a high scientific standard. acknowledgements this study is part of a project on inge lehmann and the history of modern seismology 1925–1970, funded by the carlsberg foundation and danish national archives. the author wishes to thank trine dahljensen, tine b. larsen and peter voss for help and access to material kept at geus. references bulletin of the seismological station, ivigtut 1929. copenhagen: geodetic institute, http://seis.geus.net den danske gradmaaling 1928: aarsberetning 1. april 1927 – 31. marts 1928. seismisk arbejde, http://seis.geus.net danish national archives (1): niels erik nørlund (1885–1981): mat. vedr. union géodésique et géophysique int. (1924–1936). box 116. danish national archives (2): krigsministeriet 4. kontor, indkomne sager 1868–1950, 1895–1950. geodætisk institut 50.2. box a373. danish national archives (3): niels erik nørlund (1885–1981): regnskabet 1933/1934, regnskabsmateriale vedr. carlsbergfondet (1925– 1955). box 52. danish national archives (4): generalstabens topografisk afdeling (1842–1928). vedr. etablering og drift af seismisk station ivigtut, 1924–1953. sagsarkiv (1801–1978). box 50. danish national archives (5): geodætisk institut. journalsager (1925– 1988): 3 1 udenrigsministeriet 1925 – 3 1 udenrigsministeriet 1961. box 1. doel, r.e., harper, k.c. & heymann, m. 2016: exploring greenland’s secrets: science, technology. diplomacy, and cold war planning in global context, 1–22. in: doel, r.e., harper, k.c. & heymann, m. (eds): exploring greenland. palgrave studies in the history of science and technology, new york: palgrave macmillan. geodætisk institut 1930: the seismological stations københavn and scoresbysund copenhagen: geodetic institute. geodætisk institut. 1978: geodætisk institut 1928–1978. copenhagen: geodetic institute. harboe, e.g. 1911: das erdbebenobservatorium auf der disko-insel. leipzig: wilhelm engelmann, http://seis.geus.net kragh, h., kjærgaard p.c., nielsen n. & nielsen k.h. 2008: science in denmark, a thousand-year history. aarhus: aarhus university press. lehmann, i. 1936: p’. publications du bureau central seismologique international. serie a: travaux scientifique 14, 3–31. ries, c.j. 2012: armchairs, dogsleds, ships, and airplanes. field access, scientific credibility, and geological mapping in northern and north-eastern greenland 1900–1939. in: harbsmeier, m. et.al. (eds): scholars in the field. studies in the history of fieldwork and expeditions, 329–361. aarhus: aarhus university press. author’s addresses a.l.l.j., danish national archives, rigsdagsgården 9, dk-1218 copenhagen k, denmark and geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: llj@sa.dk. geological survey of denmark and greenland bulletin 35, 2016, 17-22 17© 2016 geus. geological survey of denmark and greenland bulletin 35, 17–22. open access: www.geus.dk/publications/bull the danish pesticide leaching assessment programme (plap) was initiated in 1998 by the danish parliament in order to evaluate whether the use of approved pesticides will result in an unacceptable contamination of the groundwater, if applied under field conditions in accordance with current danish regulation. in this programme, water samples from variably saturated soil and groundwater collected at five cultivated fields are analysed for selected pesticides and their degradation products. the plap results are summarised and evaluated in yearly reports and used by the danish environmental protection agency in the regulation of pesticides in denmark (brüsch et al. 2015). in order to represent typical farming scenarios in denmark, the test fields are situated on meltwater and marine sands, and on tile-drained clayey soils in till areas. methods the five cultivated plap fields (1.2–2.4 ha), representing different soils and hydrogeological settings, spread across denmark (fig. 1) are located at silstrup, estrup and faardrup with tile-drained clayey soils, and at tylstrup and jyndevad with sandy soils (lindhardt et al. 2001). the groundwater table is shallow at all fields, which enables a rapid detection of any pesticide leaching to the groundwater (table 1). the plap fields are farmed according to conventional agricultural practice, and pesticides are applied in the maximum permissible doses and as specified in the regulations. water samples are collected weekly from drainage at the clayey till fields, and monthly from standard teflon suction cups in the unsaturated zone at the sandy fields, and from horizontal and vertical groundwater monitoring wells at all fields. the wells are installed in buffer zones surrounding the fields in order to avoid artificial transport pathways for pesticides and their degradation products from the surface to the groundwater. the vertical wells are located downstream from the field (fig. 2), except for one upstream vertical well, which is used to determine the upstream influx to the groundwater beneath the field. the horizontal wells are installed at the clayey till fields at depths of 2–3.5 m under the pesticide-treated areas, and at the sandy fields just beneath the fluctuating groundwater table. detection of pesticides or their degradation products can be directly related to the specific pesticide application to the plap fields by monitoring both the variably and fully saturated soil and accounting for potential upstream influx. in the drainage from the clayey till fields, the weighted average concentration of pesticides is based on flow-proportional sampling. in the two sandy soils, the weighted average pesticide concentration leached to the suction cups at 1 m depth is estimated from the detected concentrations and estimated percolation on a monthly basis (brüsch et al. 2015). the analytical programme includes relevant pesticides and their degradation products as well as inorganic compounds such as chloride, nitrate, phosphate and bromide, monitoring of pesticide leaching from cultivated fields in denmark walter brüsch, annette e. rosenbom, nora badawi and preben olsen silstrup estrup jyndevad fårdrup annual net precipitation < 150 mm 150–200 mm 200–250 mm 250–300 mm 300–350 mm 350–400 mm > 400 mm 100 km tylstrupsilstrup estrup jyndevad tylstrup faardrup estrup jyndevad faardrup clayey till field sandy field fig. 1. annual net-precipitation in denmark and the location of the five plap fields (http://www2.mst.dk/udgiv/publikationer/1992/87-503 -9581-5/pdf/87-503-9581-5.pdf; rosenbom et al. 2015). tylstrup and jyndevad are located in sandy areas with marine sand and glaciofluvial sand, respectively. silstrup, estrup and faardrup are situated in areas dominated by clayey till, and the three fields are drained. the sediments were deposited during and after the last glaciation. 1818 which is used as a tracer. the pesticides are generally analysed for two years following application, but the monitoring continues if significant leaching occurs. to evaluate the pesticide leaching, the water balance, including the percolation through the variably saturated soil, is assessed for all five plap fields using the numerical model macro (larsbo et al. 2005) based on long-term detailed monitoring of climate, crop-growth, soil water content, groundwater table, and if present, drainage flow (rosenbom et al. 2015). monitoring results according to the legislation of the european union, the maximum permissible concentration of any pesticide in groundwater is 0.1 μg/l (council of the european union 1994). this limit is not based on health investigations but was the analytical detection limit when the legislation was made in the 1980s, and was chosen to ensure that drinking water did not contain measurable amounts of pesticides. during the latest monitoring period from july 2012 to june 2014, a total of 7378 single analyses of different pesticides or their degradation products were carried out on water samples collected at the five sites. the leaching risk of 22 pesticides and 17 degradation products was evaluated after applying the specific pesticide on specific crops. of these 39 pesticides and their degradation products, 21 were not detected in any of the water samples. during the entire monitoring period from may 1999 to june 2014, 51 pesticides and 52 degradation products were analysed. these are listed in the appendix. the monitoring data showed leaching of 17 of the applied pesticides and their degradation products through the soil to tile drains or suction cups in average concentrations exceeding 0.1 μg/l. these are marked with asterisks in the appendix. the results of the monitoring also showed leaching of an additional 17 pesticides, but in low concentrations, marked by † in the appendix. although the concentrations exceeded 0.1 μg/l in several water samples collected from suction cups and tile drains at 1 m depth, the average leaching concentrations did not exceed 0.1 μg/l on an annual basis. in groundwater samples, twenty-one pesticides or their degradation products were only detected at concentrations table 1. characteristics of the five pesticide leaching assessment fields precipitation (mm/y)* 668 858 866 862 558 potential evapotranspiration (mm/y)* 552 555 564 543 585 area (ha) 1.1 2.4 1.7 1.3 2.3 tile drain no no yes yes yes depth to tile drain (m) 1.1 1.1 1.2 deposited by saltwater meltwater glacier glacier glacier sediment type fine sand coarse sand clay till clay till clay till topsoil classification loamy sand sand sandy clay loam sandy loam sandy loam tylstrup jyndevad silstrup estrup faardrup * based on the period 1961–1990, modified from lindhardt et al. (2001). 50 m vertical monitoring screen tile drain, inside field tile drain, outside field collector pipe groundwater flow sample point, horizontal screens horizontal monitoring screen, 3.5 m depth horizontal monitoring screen, 2 m depth 0 50 m10 m #1 outlet suction cup rain gauge buffer zone farmed area piezometer fig. 2. overview of the silstrup field and its technical installations. 19 below 0.1μg/l or not at all. these are marked by § in the appendix. at the three clayey till fields, several pesticides were detected in the drainage, whereas the frequency of detection in the groundwater monitoring screens beneath the tile drain system was lower and varied considerably between the three fields. in the two sandy fields, fewer pesticides and degradation products were generally detected, both in the variably saturated soil and in groundwater (table 2). the different leaching patterns in the sandy and clayey till fields can be attributed to specific hydrological, geological and geochemical conditions. the subsoil c horizon beneath the tile drains at the estrup field shows low permeability with few macropores (kjær et al. 2005; rosenbom et al. 2015) in contrast to the faardrup and silstrup fields, where the clayey till is characterised by fractures and heterogeneity. hence the fewer records of pesticides and degradation products in the groundwater at estrup than at faardrup and silstrup can be related to the low permeability at the former site. a comparison between the clayey till fields shows that the number of water samples containing pesticides and degradation products was higher at silstrup and estrup (35 and 40%, respectively) than at faardrup (15%). this can be attributed to different hydro-geochemical conditions and the low net precipitation at faardrup. the leaching pattern for non-pesticides shows that the average concentration of nitrate-n was much higher in both groundwater and drainage at faardrup than at the other two fields (table 2; ernstsen et al. 2015). however, the average chloride content in both drainage and groundwater at faardrup was higher than at silstrup (table 2), due to an up-concentration in the infiltration water caused by the low precipitation at faardrup. the occurrence of precipitation and subsequent percolation within the first month after application were generally higher at silstrup and estrup than at faardrup (table 1). at the clayey till fields, 59–78% of the different applied pesticides and their degradation products were detected in drainage water or groundwater (table 2), while only 28–33% of them were detected at the sandy fields. high pesticide concentrations dominated at the three clayey till fields, with 33–53% of the detections exceeding 0.1 μg/l, while only 11–16% of the detections at the two sandy fields exceeded the threshold limit. however, the limit of 0.1 μg/l is only relevant for groundwater and not for drainage water. the average nitrate concentrations were high in the groundwater of the sandy fields and lower at the clayey till fields (ernstsen et al. 2015). however, a high average nitrate concentration was recorded in both the drainage and groundwater from the faardrup field where the precipitation is low. this is probably because the uppermost part of the till is characterised by high permeability. it is therefore apparent that the pesticide and nitrate concentrations both reflect the geochemical conditions of groundwater and drainage water. further details regarding plap can be found in kjær et al. (2002, 2003, 2004, 2005, 2007, 2008, 2009, 2011), rosenbom et al. (2010), brüsch et al. (2013a, 2013b, 2015), ernstsen et al. (2015) and rosenbom et al. (2015). for further information please visit: http://pesticidvarsling.dk/monitor_uk/index.html. conclusions the results presented here provide an overall picture of the detections of pesticides and their degradation products in soil and groundwater in five monitored cultivated fields representing typical danish farming activities on clayey and sandy soils in the period from 1999 to 2014. the overall table 2. total number of pesticides analysed, detected, and detected below 0.1 μg/l in all sample types pesticides detections 16 19 39 45 38 and detections >0.1 μg/l 6 9 22 31 21 metabolites detections in % 28.1 32.8 59.1 77.6 66.7 >0.1 μg/l in % 10.5 15.5 33.3 53.4 36.8 groundwater avg nitrate-n 15.5 11.9 3.0 0.4 8.5 chloride 49.9 15.6 29.5 11.7 27.1 drainage avg nitrate-n ns ns 2.1 3.5 11.2 chloride ns ns 30.3 26.6 27.5 fine-grained sand coarse-grained sand clayey till tylstrup jyndevad silstrup estrup faardrup samples collected from suction cups, drainage and groundwater in the five plap fields between 01 january 2000 and july 2012. average nitrate and chloride concentrations from groundwater and drainage in the period january 2011 – july 2012. avg: average concentration in mg/l. ns: no samples. 2020 pesticide leaching detected in the monitoring programme is an outcome of the pesticide selection, hydraulic conditions, type of agriculture and the geochemical conditions such as the redox potential, aerobic conditions and hence the leaching of nitrate-n and potential persistence of individual pesticides. for instance, the leaching of pesticides is more pronounced in fractured clayey soils than in sandy soils due to fast transport in anaerobic fractures in the former soils, in contrast to slower matrix transport in the more aerated sandy soils. this is illustrated by the high number of recorded pesticides in drainage water and groundwater from clayey till soils due to bypassing of the topsoil by rapid leaching through well-connected macropores such as wormholes and fractures (rosenbom et al. 2015). the occurrence of pesticides in samples from the two sandy soils is probably specifically linked to the application of persistent pesticides such as metalaxyl-m applied to potatoes. references brüsch, w., kjær, j., rosenbom, a.e., juhler, r.k., gudmundsson, l., plauborg, f., nielsen, c.b. & olsen, p. 2013a: the danish pesticide leaching assessment programme: monitoring results may 1999 – june 2011, 108 pp. copenhagen, denmark: geological survey of denmark and greenland. brüsch, w., rosenbom, a.e., juhler, r.k., gudmundsson, l., plauborg, f., nielsen, c.b. & olsen, p. 2013b: the danish pesticide leaching assessment programme: monitoring results may 1999 – june 2012, 106 pp. copenhagen, denmark: geological survey of denmark and greenland. brüsch, w., rosenbom, a.e., badawi, n., v. platten-hallermund, f., gudmundsson, l., plauborg, f., nielsen, c.b., laier, t. & olsen, p. 2015: the danish pesticide leaching assessment programme: monitoring results may 1999 – june 2013, 110 pp. copenhagen, denmark: geological survey of denmark and greenland. council of the european union 1994: council directive 94/43/ec establishing annex vi to directive 91/414/eec concerning the placing of plant protection products on the market. official journal of the european union l227, 1.9.1994, 31–55. ernstsen, v., olsen, p. & rosenbom, a.e. 2015: long-term monitoring of nitrate transport to drainage from three agricultural clayey till fields. hydrology and earth system sciences 19, 3475–3488, http:// dx.doi.org/10.5194/hess-19-3475-2015. kjær, j. et al. 2002: the danish pesticide leaching assessment programme: monitoring results may 1999 – june 2001, 150 pp. copenhagen, denmark: geological survey of denmark and greenland. kjær, j., ullum, m., olsen, p., sjelborg, p., helweg, a., mogensen, b., plauborg, f., grant, r., fomsgaard, i. & brüsch, w. 2003: the danish pesticide leaching assessment programme: monitoring results may 1999 – june 2002, 158 pp. copenhagen, denmark: geological survey of denmark and greenland. kjær, j., olsen, p., barlebo, h.c., juhler, r.k., plauborg, f., grant, r., gudmundsson, l. & brüsch, w. 2004: the danish pesticide leaching assessment programme: monitoring results may 1999 – june 2003, 146 pp. copenhagen, denmark: geological survey of denmark and greenland. kjær, j., olsen, p., barlebo, h.c., juhler, r.k., henriksen, t., plauborg, f., grant, r., nyegaard p. & gudmundsson, l. 2005: the danish pesticide leaching assessment programme: monitoring results may 1999 – june 2004, 86 pp. copenhagen, denmark: geological survey of denmark and greenland. kjær, j., olsen, p., barlebo, h.c., henriksen t., plauborg, f., grant, r., nyegaard, p., gudmundsson, l. & rosenbom, a.e. 2007: the danish pesticide leaching assessment programme: monitoring results may 1999 – june 2006, 99 pp. copenhagen, denmark: geological survey of denmark and greenland. kjær, j., rosenbom, a., olsen, p., juhler, r.k., plauborg, f., grant, r., nyegaard, p., gudmundsson, l. & brüsch, w. 2008: the danish pesticide leaching assessment programme: monitoring results may 1999 – june 2007, 91 pp. copenhagen, denmark: geological survey of denmark and greenland. kjær, j., rosenbom, a., olsen, p., ernstsen, v., plauborg, f., grant, r., nyegaard, p, gudmundsson, l. & brüsch, w. 2009: the danish pesticide leaching assessment programme: monitoring results may 1999 – june 2008, 88 pp. copenhagen, denmark: geological survey of denmark and greenland. kjær, j., rosenbom, a.e., olsen, p., ernstsen, v., plauborg, f., grant, r., gudmundsson, l. & brüsch, w. 2011: the danish pesticide leaching assessment programme: monitoring results may 1999 – june 2010, 110 pp. copenhagen, denmark: geological survey of denmark and greenland. larsbo, m., roulier, s., stenemo, f., kasteel, r. & jarvis, n. 2005: an improved dual-permeability model of water flow and solute transport in the vadose zone. vadose zone journal 4, 398–406. lindhardt, b., abildtrup, c., vosgerau, h., olsen, p., torp, s., iversen, b.v., jørgensen, j.o., plauborg, f., rasmussen, p. & gravesen, p. 2001: the danish pesticide leaching assessment programme: site characterization and monitoring design, 73 pp. copenhagen, denmark: geological survey of denmark and greenland. rosenbom, a.e., brüsch, w., juhler, r.k., ernstsen, v., gudmundsson, l., plauborg, f., grant, r. & olsen, p. 2010: the danish pesticide leaching assessment programme: monitoring results may 1999 – june 2009, 102 pp. copenhagen, denmark: geological survey of denmark and greenland. rosenbom, a.e., olsen, p., plauborg, f., grant, r., juhler, r.k., brusch, w. & kjaer, j. 2015: pesticide leaching through sandy and loamy fields – long-term lessons learnt from the danish pesticide leaching assessment programme. environmental pollution 201, 75–90. authors’ addresses w.b.*, a.e.r. & n.b., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. p.o., department of agroecology, aarhus university, blichers allé 20, dk-8830 tjele, denmark. *present address: danmarks naturfredsningsforening, madsnedøgade 20, dk-2100 copenhagen ø, denmark. e-mail: wb@dn.dk 21 appendix. plap analyses from may 1999 to june 2014. part a aclonifen aclonifen † 111 § 298 amidosulfuron amidosulfuron 144 3 1 0.11 § 332 desmethyl-amidosulfuron 24 88 aminopyralid aminopyralid † 133 § 261 azoxystrobin azoxystrobin * 717 139 16 1.4 1798 3 0.04 cypm * 740 390 144 2.1 1910 69 5 0.19 bentazone bentazone * 1051 350 43 43 2603 81 8 0.6 2-amino-n-isopropyl-benzamide 561 4 0.06 1295 1 0.02 bifenox bifenox 303 17 3 0.38 751 7 1 0.1 bifenox acid * 278 55 47 8.6 702 29 23 3.1 nitrofen 303 11 4 0.34 751 boscalid boscalid † 56 § 111 bromoxynil bromoxynil 528 5 3 0.6 § 1122 chlormequat chlormequat † 95 2 0.017 § 190 clomazon clomazone 224 1 1 0.28 § 598 fmc 65317 216 1 1 0.3 577 clopyralid clopyralid 219 7 4 4.094 § 520 1 0.026 cyazofamid cyazofamid † 100 § 262 desmedipham desmedipham † 287 580 1 0.033 ehpc 199 383 diflufenican diflufenican 109 32 14 0.49 324 1 1 0.47 ae-b107137 * 121 19 1 0.13 333 1 0.016 ae-05422291 109 324 dimethoate dimethoate † 515 1 1 1.417 1253 1 0.085 epoxiconazole epoxiconazole 330 14 2 0.39 999 1 0.011 ethofumesate ethofumesate * 519 70 17 12 1095 36 7 1.4 fenpropimorph fenpropimorph † 657 2 0.038 1531 2 0.029 fenpropimorph acid 636 2 1 0.25 1435 flamprop-m flamprop-m-isopropyl 520 38 1 0.109 1204 1 0.024 flamprop 525 23 1 0.35 1212 florasulam florasulam † 146 351 florasulam-desmethyl 109 130 fluazifop-p-buthyl fluazifop-p-butyl 128 232 tfmp * 184 53 24 0.64 555 87 16 0.29 fluazifop-p 451 11 4 3.8 1109 7 1 0.17 fludioxonil cga 192155 † 11 § 48 cga 339833 11 48 fluroxypyr fluroxypyr * 521 4 3 1.4 1273 2 0.072 glyphosate glyphosate * 1091 429 136 31 2216 77 5 0.67 ampa * 1092 632 142 5.4 2217 37 0.08 iodosulfuron-methyl-natrium metsulfuron-methyl 332 1 0.054 842 iodosulfuron-methyl † 60 § 250 ioxynil ioxynil 527 24 7 0.25 1128 1 0.01 linuron linuron † 67 § 271 mancozeb etu 44 7 0.038 200 2 0.024 ebis 7 25 mcpa mcpa 354 14 3 3.894 916 1 0.019 2-methyl-4-chlorophenol 354 2 1 0.24 912 mesosulfuron-isopropyl mesosulfuron-methyl 153 13 0.059 § 411 mesosulfuron 119 119 mesotrione mesotrione † 50 § 156 amba 50 156 mnba 50 156 pesticide analyte tile drain and suction cup groundwater samples det. ≥0.1 max. samples det. ≥0.1 max. fifty-one pesticides and 52 degradation products analysed in the plap programme in the period may 1999 – june 2014. the columns show the number of water samples analysed, number of detections, and detections in concentrations ≥ 0.1μg/l in water samples from the variably-saturated zone (drainage and suction cups), and in groundwater (vertical and horizontal groundwater wells). det: number of detections. ≥0.1: number of detections ≥0.1μg/l. max: maximum concentration in μg/l. *: pesticides and their degradation products leached through soil to tile drains or suction cups in average concentrations above 0.1 μg/l. †: pesticides not detected or detected only in a few samples above their threshold concentrations at 1 m depth. §: pesticides and their degradation products not detected or only detected in a few samples in groundwater. 2222 appendix. plap analyses from may 1999 to june 2014. part b metalaxyl-m metalaxyl-m 207 15 0.037 592 79 23 1.3 cga 108906 * 215 175 69 4.8 593 468 128 2.7 cga 62826 * 216 100 25 1.2 593 147 8 0.68 metamitron metamitron * 515 103 31 26.369 1095 53 7 0.63 desamino-metamitron * 518 129 23 5.549 1094 78 16 1.3 metrafenone metrafenone 136 20 0.072 273 1 0.04 metribuzin metribuzin 97 2 0.024 414 1 0.014 diketo-metribuzin 340 256 63 0.69 552 479 336 1.372 desamino-diketo-metribuzin * 255 81 51 2.1 551 256 18 1.831 desamino-metribuzin * 91 392 pendimethalin pendimethalin 694 89 30 32 1811 1 0.052 phenmedipham phenmedipham 288 580 2 0.025 mhpc 288 2 1 0.19 580 1 0.053 3-aminophenol 109 245 picolinafen picolinafen 117 18 0.07 193 cl153815 * 117 31 11 0.5 193 pirimicarb pirimicarb 887 62 0.077 2120 6 0.035 pirimicarb-desmethyl-formamido * 707 29 13 0.379 1638 2 0.076 pirimicarb-desmethyl 780 8 0.053 1911 3 0.042 propiconazol propiconazole 899 32 3 0.862 2084 3 0.035 propyzamid propyzamide * 257 27 8 1.6 754 10 2 0.14 rh-24644 257 19 0.051 754 2 0.032 rh-24580 257 2 0.016 754 rh-24655 233 1 0.017 690 prosulfocarb prosulfocarb 199 6 1 0.18 516 5 0.032 pyridat pyridate 39 116 phcp 125 4 4 2.69 373 14 4 0.309 rimsulfuron rimsulfuron 117 367 ppu * 502 388 74 0.29 1519 432 13 0.23 ppu-desamino 502 186 6 0.18 1519 107 0.089 tebuconazole tebuconazole * 289 47 17 2 784 8 2 0.12 1.2.4-triazol * 16 7 1 0.17 terbuthylazine terbuthylazine * 513 213 56 11 1324 88 23 1.9 desethyl-terbuthylazine * 612 365 88 8.3 1664 261 33 0.94 desisopropylatrazine 414 156 2 0.44 996 92 0.047 hydroxy-terbuthylazine * 384 136 18 0.99 940 34 0.069 2-hydroxy-desethyl-terbuthylazine * 342 128 28 6.3 850 9 0.092 thiacloprid thiacloprid † 47 § 100 thiacloprid-amide 47 1 0.012 100 m34 55 100 thiacloprid sulfonic acid 56 100 thiamethoxam thiamethoxam † 132 § 359 cga 322704 132 359 triasulfuron triasulfuron † 82 § 301 triazinamin 393 1103 1 0.042 tribenuron-methyl triazinamin-methyl † 569 2 0.042 § 1523 triflusulfuron-methyl triflusulfuron-methyl † 95 288 in-e7710 95 5 0.014 288 in-m7222 95 288 1 0.052 in-d8526 95 288 pesticide analyte tile drain and suction cup groundwater fifty-one pesticides and 52 degradation products analysed in the plap programme in the period may 1999 – june 2014. the columns show the number of water samples analysed, number of detections, and detections in concentrations ≥ 0.1μg/l in water samples from the variably-saturated zone (drainage and suction cups), and in groundwater (vertical and horizontal groundwater wells). det: number of detections. ≥0.1: number of detections ≥0.1μg/l. max: maximum concentration in μg/l. *: pesticides and their degradation products leached through soil to tile drains or suction cups in average concentrations above 0.1 μg/l. †: pesticides not detected or detected only in a few samples above their threshold concentrations at 1 m depth. §: pesticides and their degradation products not detected or only detected in a few samples in groundwater. samples det. ≥0.1 max. samples det. ≥0.1 max. geological survey of denmark and greenland bulletin 31, 2014, 9-14 9 125 years of geological research for society johnny fredericia and peter gravesen in 1888 the first geological survey in the kingdom of denmark was born as the geological survey of denmark (dgu, danmarks geologiske undersøgelse) and in 1946 the geological survey of greenland (ggu, grønlands geologiske undersøgelse) was established. both surveys were located in copenhagen and were amalgamated in 1995 to form the geological survey of denmark and greenland (geus). this was a happy amalgamation for both surveys. even though dgu and ggu had different backgrounds and different working areas the employees had a lot in common, and today geus continues dgu’s and ggu’s work, but in an integrated way with new synergies – so a stronger survey has evolved. this was the reason for celebrating geus’ 125th anniversary on 4 april 2013 and it also explains the logo with the years 1888 and 1946 to the left and 2013 to the right (fig. 1). the event was celebrated with a series of lectures held by national and international geoscientists, a reception for invited guests and a gala night for geus staff and board. the geological survey of denmark 1888–1995: geology for society the document shown in fig. 2 is considered the birth certificate of the first geological survey in the kingdom of denmark: the geological survey of denmark (dgu). the letter, dated 4 april 1888, grants the first danish professor of geology j.f. johnstrup from the mineralogical museum government funds to organise and start the geological mapping of denmark. it was estimated to take 20 years and cost 20 000 danish kroner per year – corresponding to a little less than 1 million euro per year in today’s money. establishing a national geological survey was not an entirely new idea. denmark lagged far behind its neighbours, who had started more than 25 years earlier – and it was one of the convincing arguments in parliament at the time. the geological surveys of norway and of sweden were both established in 1858 (fig. 3). the survey programmes in the © 2014 geus. geological survey of denmark and greenland bulletin 31, 9–14. open access: www.geus.dk/publications/bull fig. 1. geus’ 125th anniversary logo. fig. 2. the birth certificate of the geological survey of denmark from 1888 granting funds to start geological mapping of denmark. 1010 other nordic countries were also much more extensive than the modest beginnings in denmark. the purpose of the new survey was to map the superficial deposits of the country to obtain knowledge about the important geological resources. at that time it was largely gravel, sand, clay, peat, brown coal and limestone, but also hydrogeology and the quality of agricultural soils were important. knowledge of the superficial deposits was also of strategic importance for military operations. figure 4 shows the first published map and one of the most recent. the same mapping units have been used throughout the years, which means that it is now possible to produce seamless digital maps. mapping and investigating the quaternary deposits in denmark were core tasks for the survey for more than a century, but over the years the systematic mapping of the superficial deposits has been overshadowed by large numbers of other tasks and commitments. a few examples are mentioned below. danish brown coal was previously widely used for heating, and exploration of brown-coal deposists in jylland during world war i and up to world war ii was a major task for dgu. the mapping of marl deposits was another important task for the young survey; local marl was used to improve the agricultural soils in denmark. investigation of palaeoclimate and vegetation history based on macrofossils and pol1888 1858 1858 1835 1845 1849 1859 1873 1852 1896 1860 1867 1885 1882 1919 1919 1940 1937 1936 1936 c. 1930 1865 188218671949 1909 1946 fig. 3. the years of establishment of the european geological surveys (påsse & kim-andersson 2008). geology of the pre-quaternary surface nedre miocæn, ler og sand lower miocene, clay and sand oligocæn, ler oligocene, clay eocæn, moler (diatomit med askelag) eocene, clayey diatomite with ash layers palæocæn, plastisk ler paleocene, clay danien, kalk danian, limestone øvre kridt, skrivekridt upper cretaceous, chalk prækvartæroverfladens højdekurver i meter contours in meters of the pre-quaternary surface normalforkastning normal fault glacialtektonisk forkastning glacial tectonic thrust fault 0 5 km 0 -25 -50 -75 e f 6310000 56°55'0 56°50'0 6300000 56°45'0 6290000 thisted bredning limfjorden visby bredning dragstrup vig vil su nd thisted nykøbing m kås bredning sal ling su nd livø bredning assels hage søndervig lindholm stenklipperne a g e r ø dover mølle grund karby vig hage gudnæs skyum øre erslev kærgl. jølby frøslev vang sindbjerg e f c d flade sejerslev ejerslev 0 0 0 0 0 0 0 -25 -25 -25 -25 0 0 -25 -25 -50 -25 -25 -25 -25 -25 -25 -25 -25 -25 -25 -100 -75 -75 -75 -50 -25 -25 -50 -50 -25 -25 -50 +25 -75 -50 -50 -25 -25 +25 +25 0 0 0 0 -25 0 0 0 0 0 -50 -50 -50 -50 -75 -50 0 -25 -50 dc visby bredning damsgård mollerup frøslevvang nykøbing mors thisted saltdome erslev saltdiapir denmark 50 km denmark 50 km a b fig. 4. old and recent published maps of deposits at the terrain surface. the deposits are mainly of quaternary age. a: map covering parts of north-east sjælland, published on a scale of 1:40 000 (rørdam 1893). b: one of the latest maps covers the island of mors in north-west jylland, published on a scale of 1:50 000 (pedersen & jakobsen 2012). 11 len from bog and lake sediments was another research area that started early at dgu. danish water supply has been increasingly based on groundwater, and at an early stage dgu became engaged in solving quantity and quality problems for the public. in the 1970s and the 1980s the survey was involved in groundwater mapping, and during this period the first large digital database in denmark was established with borehole and groundwater data. several projects concerning pollution of the groundwater with nitrate and pesticides were initiated in the 1980s and 1990s. special attention was put into locating areas in denmark which were sensitive to pesticide pollution and the danish pesticide leaching assessment programme was established. further nation-wide groundwater mapping was carried out in the early 2000s. interest in oil and gas exploration in denmark began in 1935, and dgu became involved in aspects of this work, analysing borehole samples and interpreting seismic data. this work also resulted in the production of a series of thematic geological maps. the samba database was established to store borehole data, geophysical surveys and reports. in the 1970s when it was suggested that radioactive waste from potential nuclear power plants in denmark would be stored in permian salt deposits, dgu demonstrated that salt diapirs were not suitable repositories. this was one of the reasons why denmark decided in 1985 not to build nuclear power plants to produce electricity. from 1888 to 1987 dgu was also the geological survey of the faeroe islands and mapped the volcanic rocks and superficial deposits of the islands. in addition, dgu carried out initial investigations to map potential oil reservoirs and conducted the first exploration boreholes. the geological survey of greenland 1946–1995: a survey at the frontier after k.l. giesecke’s initial studies during the napoleonic wars, other geological investigations and mineral exploration in greenland began around 1850. from 1879 the expeditions were organised under the auspices of the commission for scientific investigations in greenland, often covering a wide range of natural science topics. geological investigations later became a key element in a number of expeditions to east greenland under the leadership of lauge koch, funded by the danish state. the expeditions began in 1926 and continued until the outbreak of world war ii. they were resumed after the war and lasted until 1958. in west greenland geological mapping began in 1946 following the establishment of the geological survey of greenland (ggu). the systematic mapping of greenland was a tremendous task; greenland is characterised by very limited infrastructure, and the size of the ice-free area is c. 410 000 km2, nearly the size of sweden. as the knowledge of greenland geology was initially very limited, the first years were spent on a reconnaissance study of the west coast to gain an impression of the geology involved. systematic mapping on a scale of 1:100 000 began in the ivittuut area in 1956. ten years later it was realised that on that scale it would take a lifetime to cover greenland, so the mapping was augmented with a series on 1:500 000 scale, first in west greenland, then in east greenland and in north greenland from the 1970s (figs 5, 6). the systematic 10 km 63°ngreenland 51°w fig. 5. excerpt from the seamless geological map of greenland, published on a scale of 1:500 000. 1212 mapping on 1:500 000 scale was completed in 40 years from 1964 to 2004. study and evaluation of mineral deposits of potential economic interest also took place along with the geological mapping. this was augmented in later years with airborne geophysical surveys with focus on mineral deposits. these studies, including geophysical investigations of offshore sedimentary basins, still continue, most recently in 2013 in south-east greenland. in the 1960s, when denmark was considering use of nuclear power, survey emphasis was placed on the discovery of radioactive minerals, as the danish state had excluded radioactive minerals from commercial prospecting licences. kvanefjeld in south greenland was identified as a target in the 1960s and was extensively drilled (fig. 7). later stream sediment surveys throughout greenland have helped to locate gold and diamond provinces and indicate potential environments for base metals, platinum-group elements and rare-earth elements. following the discovery of oil in alaska there was interest in the sedimentary basins offshore west greenland and six licence blocks were issued in 1974. this entailed extensive detailed studies by the survey of the adjacent onshore sedimentary areas. it was followed by active airborne and seaborne geophysical surveys and further study of onshore sedimentary basins in both west and east greenland. greenland is dominated by a large ice sheet that influences the climate of the northern hemisphere. the run-off from this ice sheet was investigated in west greenland prior to the building of hydroelectric plants, and the advances and retreats of its outlet glaciers have been followed and reconstructed backwards for decades. with the advent of climatic change the study of mass-balance models for the whole ice sheet has intensified, and in recent years geus has established a network of automatic weather stations which allows detailed monitoring of changes in the mass balance of the ice sheet (promice; fig. 8). the geological survey of denmark and greenland – the first challenges a reorganisation of the raw materials administration between denmark and greenland took place in 1995. as part of this reorganisation the future of ggu was discussed. the two directors, martin ghisler from ggu and ole winther christensen from dgu, had the same vision and worked to amalgamate ggu and dgu and make a larger and stronger geological survey (fig. 9). cooperation with geologists form the university of copenhagen was strengthened by the estabfig. 6. base camp in warming land in north greenland in 1985. an islandic twin otter aircraft and two swiss helicopters are parked at the tent camp. the 600 m high cliffs in the background consist of palaeozoic platform carbonates. photograph: jakob lautrup. fig. 7. ten kilometres of core were drilled between 1958 and 1981 at kvanefjeld in south greenland to investigate the uranium contents. the drilling teams and geologists were camped. photograph: jan bondam. fig. 8. setting up an automatic weather station near the margin of the greenland ice sheet. photograph: claus heinberg. 13 lishment of geocenter copenhagen that comprised geus, the geological institute, the geographical institute and the geological museum of the university of copenhagen. in 2007 an act for geus was passed by the danish parliament, and geocenter denmark was created, with a formalised collaboration between geus, the university of copenhagen and the university of aarhus. when geus celebrated its 125th anniversary, it was a celebration both for the survey and for a strong alliance with the integrated danish geoscience community. dgu and ggu had different beginnings and initially different schedules like trains meeting at a station, but we have stayed together and share the full story. therefore it was logical to celebrate 125 years. a book describing the history of geus: ‘we – the people down-to-earth’ (vi de jordbundne) was published by geus in connection with the anniversary (hansen & thomsen 2013). its subtitle: ‘glimpses of geus over 125 years’ (glimt af geus gennem 125 år). the book focuses on geus’ role in society through the years and is partly based on earlier publications (rasmussen 1988; ellitsgaard-rasmussen 1996; fig. 10). the book does not give the complete story of geus, but provides an overview and glimpses of its history. it comprises special highlights, important changes of the institution and tells about people at geus in the past and at present. the book also includes a number of anecdotes that capture the work and social culture at geus. geus today geus is striving to be an international research organisation with the purpose of advising its users, the surrounding society in a broad sense. geus also develops and maintains national databanks with geological information on water, energy and mineral deposits and to some degree the climate, and is thus the geological memory of danish society. being increasingly internationally orientated, geus has many partners in the nordic countries, in the rest of europe and farther afield. in particular, geus shares the vision of a much closer collaboration between the geological surveys of europe, e.g. eurogeosurveys. a new regional cooperation between the surveys around the north atlantic (nag, northeast atlantic geosciences) and increased cooperation between the nordic countries have created important new avenues to develop our common role as surveys for a broader public. however, other research organisations are also important partners for geus. for instance, geus has been partner in more than 75 european union projects. geus also has a number of memoranda of understanding with countries outside the eu, and has been working in more than 30 countries outside europe over the past 10 years. for example, geus has worked together with vietnam for 17 years to build the country’s capacity to develop their oil, gas and water resources. geus also forms partnerships with other users of our knowledge, where we can contribute to growth in denmark and greenland. water technology as a new business adventure is currently high on the political agenda, and geus strongly supports this effort. it is the intention to strenghten geus’ unique cooperation with copenhagen and aarhus universities in geocenter denmark, since it is a cornerstone in the future development of geus. geus in the future geus’ goals for the future are expressed in the institution’s strategy: geus strategi 2012. it stresses that geus fig. 9. dgu and ggu amalgamated in 1995. left: the headquarters of geus at øster voldgade 10 in copenhagen. right: from the left, the former directors martin ghisler and ole winther christensen in discussion with niels henriksen and jens morten hansen on a field trip to north-east greenland in 1995. photographs: jakob lautrup. 1414 will focus on increased cooperation and partnership – here and abroad, based on geus’ nine strategic topics shown in fig. 11. the strategic topics coincide with several of the grand challenges of this century: minerals, environment and climate change, water resources, energy supply and energy storage. these challenges match geus’ own ambitions, and geus is ready to deal with the challenges where geoscience plays an important role. this reflects that geus is responsible and constantly trying to adjust itself to the changing needs of our society, which is becoming more and more internationally integrated. geological knowledge is of crucial importance to society in the 21st century. geus will follow the changing demands from society and focus on the geological knowledge required so this can be developed to the benefit of danish and international societies. references ellitsgaard-rasmussen, k. 1996: en stjerne fødes. beretning om ggu’s tilblivelse, 76 pp. danmarks og grønlands geologiske undersøgelse rapport 1996/102, 76 pp. geus 2012: geus strategy 2012, basis for performance conctract 2012–2015, 28 pp. copenhagen: geological survey of denmark and greenland. hansen, j.m. & thomsen, h.h. (eds) 2013: vi, de jordbundne. glimt af geus gennem 125 år (1888–2013), 160 pp. copenhagen: geological survey of denmark and greenland. pedersen, s.a.s. & jakobsen, p.r. 2012: geological map of denmark 1:50  000, mors. copenhagen: geological survey of denmark and greenland. påsse, t. & kim-andersson, a. (eds) 2008: sveriges geologiska undersökning. 150 år i samhällets tjänst, 185 pp. uppsala: sveriges geologiska undersökning. rasmussen, l.b. 1988: en jordisk krønike. træk af dgu’s historie 1888– 1998, 114 pp. copenhagen: danmarks geologiske undersøgelse. rørdam, k. 1893: kortbladene helsingør og hillerød. danmarks geologiske undersøgelse i. række 1, 110 pp. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: pg@geus.dk fig. 10. the front cover of the anniversary book (hansen & thomsen 2013) shows a geus drilling team in north-east greenland. photograph: jørgen bojesen-koefoed. dissemination of digital data and knowledge the strategic topics water resources under pressure oil and gas supply and transition to green energy geothermal energy and heat storage geology and health minerals – basic and critical resources geology in the public arena geology across land and sea past and future climate fig. 11. the nine strategic topics from ‘geus strategy 2012 – basis for performance contract 2012–2015’ coincide with several of the grand challenges of this century: minerals, environment and cli mate change, water resources and energy supply, and energy storage (geus 2012). mailto:hl@geus.dk e2019430206-01 global population has increased rapidly in recent decades. so far, it has been possible to feed the growing population by using more and more land for agriculture, using irrigation and artificial fertilisers and by improving the efficiency of agriculture. recently the growth of the global agricultural area has slowed. however, the need for food will continue to grow markedly in coming years. this demand can no longer be met by using increasingly more land for agriculture, and in many areas it is not possible to increase crop production by irrigation (wise 2013). large areas in the tropics are characterised by strongly depleted soils with low concentrations of nutrients such as nitrogen, phosphorous and potassium. in such areas, the yield of crop per hectare is much lower than the theoretical yield using optimal fertilising (ray et al. 2013). reducing the gap between real and potential crop productivity offers the best solution to achieve food security for the world’s rapidly growing population. poor soil quality in the tropics is largely due to the rapid weathering of minerals and leaching of dissolved nutrients in the warm and humid climate. if weathered minerals are not replaced by new minerals, for example due to volcanic activity, then soil fertility continues to decline over time. therefore, it is necessary to use increasing amounts of fertilisers to feed growing populations in the tropics. most nutrients come from geological deposits; the only exception is nitrogen, which can be extracted from the atmosphere. nutrients that are mined constitute a limited resource. hence the known occurrences of phosphorous can only cover the current demand for a few decades (van vuuren et al. 2010). in recent years, investigations have been conducted to see if the productivity of nutrient-poor soils can be improved by the application of glacial rock flour from greenland. rock flour in southern west greenland consists of fine-grained silt, formed by the grinding of bedrock by stones and boulders embedded in the basal part of glaciers. preliminary results indicate that plants cultivated in soils with rock flour can achieve increased growth (m.t. rosing, unpublished data 2019). however, the research is still in its early days and many questions remain. we do not know why adding rock flour to soil results in increased growth. maybe the silt fraction improves the soil properties. also we do not know if it is feasible to mine rock flour and transport it to the tropics. as a first step towards answering some of these questions, our aim here was to simply map and sample the glacial rock mapping glacial rock flour deposits in tasersuaq, southern west greenland ole bennike*1, jørn bo jensen1, frederik næsby sukstorf2 and minik t. rosing3 geus bulletin is free to individuals and institutions in electronic form. the author(s) retain copyright over the article contents. research article | open access geus bulletin vol 43 | e2019430206 | published online: 17 july 2019 https://doi.org/10.34194/geusb-201943-02-06 grab sample core site greenland 3 km 65°n 51°w fig. 1. satellite image of the tasersuaq area showing sampling sites. inset: the location of tasersuaq in greenland. https://doi.org/10.34194/geusb-201943-02-06 e2019430206-02 flour in tasersuaq, a large proglacial lake in southern west greenland, c. 105 km north-east of nuuk. lake tasersuaq the area surrounding lake tasersuaq (65°n, 50.8°w) is characterised by large u-shaped valleys and ice-scoured rounded mountains that reach elevations of 600 to 800 m asl. the bedrock geology is dominated by archaean basement rocks, mainly qôrqut granite (escher & pulvertaft 1995). the lake is 30 km long, 2–3 km wide and has an irregular shape (fig. 1). the elevation of the lake is 74 m above sea level, which is close to the local marine limit (i.e. the limit of raised marine deposits). however, we did not observe any marine deposits, suggesting that the lake threshold is above the marine limit and thus the deposits are free from nacl, which is toxic to many plants. a 28 km long, land-based outlet glacier, saqqap sermersua, from the greenland ice sheet almost reaches a north-eastern branch of the lake, but is separated by a sandur plain and delta. there are no fresh, unvegetated moraines or trim line zone bordering the glacier, which appears to have advanced in recent centuries (weidick et al. 2012). huge amounts of meltwater drain into tasersuaq and the lake water is highly turbid. to the south-west of the lake is a large plain, narsarsuaq, which mainly consists of glaciofluvial sand and some gravel that was deposited during the last deglaciation of the region, when lake tasersuaq was still filled by glacier ice and meltwater drained towards the south-west. several low u-shaped moraine ridges are found on the plain (weidick 1971); they mark stillstands or minor re-advances during periods of ice retreat. a number of distinct kame terraces are found at the south-eastern end of the lake, probably formed by a thinning glacier lobe in the lake basin. the terraces were mapped as ice margin deposits by weidick (1978). larsen et al. (2014) assigned the ice margin deposits in the tasersuaq area to the kapisigdlit stade – a holocene glacial stade, which they suggested has an age of 10.1–10.4 ka bp. from this and 10be exposure ages obtained from the region by larsen et al. (2014), we suggest that the tasersuaq lake basin was deglaciated shortly after 10 ka bp. a c d b 2 cm fig. 2. photographs from the field work. a: the geus ii speedboat with the streamer on the foremost part of the boat and the two c-boom floats in the water. b: deploying the coring platform on lake tasersuaq. c: the vibrocore with a grey sediment core in the tube, on the deck of the coring platform. d: layered, probably varved, silt and clay sediments retrieved from the lake floor. e2019430206-03 methods in 2018, we carried out echo sounding, seismic profiling, grab sampling and coring at lake tasersuaq. the acoustic work was conducted mainly onboard the geus ii speedboat (fig. 2a). the boat was equipped with a gps system for precise positioning and a navisound 215 echo sounder to measure water depth down to about 100 m. for seismic profiling we used a c-boom seismic source and a geo-sense ministreamer, with a frequency band of 0.5–1.5 khz. a resolution of c. 30 cm can be achieved, and the penetration was about 50 m. the streamer was kept 4 m from the propeller wash by a glass fibre rod. for the acquisition of the seismic data we used a chesapeake technology 24 bit ad converter and for run line management, recording and processing of seismic data we used chesapeake sonarwiz 6. sediment sampling was conducted from a uwitec coring platform kept afloat by four inflatable tubes (fig. 2b). the platform had a payload of 2600 kg and a 4 m high tripod and winch, which was used to pull the sediment cores up from the lake floor. we mounted a 25 hp outboard motor for transportation of the platform on the lake. a small zodiac dinghy was used for transport between the shore and the platform – and for safety as a man-over-board boat. we collected a total of 63 grab samples and seven sediment cores, noting their position with a hand-held gps. grab samples were collected using a small van veen sampler. often, we only retrieved small samples due to the stiff nature of the sediments and so we collected multiple samples at each site to get enough sediment for analysis. sediment cores were collected using a vibecore-d developed by sdi speciality devices. the corer consists of a battery-operated vibrating core head, a weight ring and a core tube and works by vibrating core tubes down into the sediments. the sediment cores are kept in the tubes by valves in the core head and by core catchers. we collected 2 m long, 7.6 cm diameter sediment cores, using polycarbonate core tubes (fig. 2c). all sediment a b >50 0–10 10–20 20–30 30–40 40–50 50–60 60–70 70–80 water depth (m) 80–90 90–100 100–120 120–177 0–10 10–20 20–30 30–40 40–50 50–60 60–70 70–80 thickness (m) 80–90 > 50 65°n 3 km 50°50´w fig. 3. a: bathymetry of lake tasersuaq determined from echo sounder and c-boom seismic profiles. b: sediment thickness. the area marked with ‘> 50’ indicates sediment thickness greater than 50 m. here, the boomer signal could not penetrate the sediment succession any farther and so the exact thickness could not be determined. e2019430206-04 samples and cores were reserved for analyses at a later date, to characterise their grain size distributions, mineralogical and elemental compositions. here, we focus only on the mapping of sediments. results and discussion the southern and northern parts of tasersuaq are relatively shallow, with water depths generally below c. 40 m, and just 5–10 m in the south-west (fig. 3a). water depth off the sandur plain and delta in front of the saqqap sermersua glacier increases rapidly, reaching 177 m in the central north–southorientated part of the lake. glacial rock flour is found in most parts of the lake, but sandy sediments dominate the delta in front of saqqap sermersua. sandy sediments were also found on the lake floor in front of the delta, at water depths up to 30 m. deposition of sand in such deep waters is probably due to strong bottom currents. sandy sediments also dominate the shallow southwestern branch of tasersuaq and we suggest that they represent glaciofluvial sediments that accumulated during the last deglaciation. a thin layer of glacial rock flour was found at some sampling sites in the south-west, but this fine-grained material is probably eroded and redeposited in deeper waters during stormy weather. 200 m 10 m 200 m 10 m a b c ssw nne nw se ice berg plough marks bedrock surface b 3 km a bedrock high bedrock surface fig. 4. two examples of seismic profiles from lake tasersuaq. a: line 0004, b: line 0013. c: location of seismic data aquired in 2018 and the location of the seismic profiles shown in a and b. *corresponding author: ole bennike | e-mail: obe@geus.dk 1 geological survey of denmark and greenland (geus), c.f. møllers allé 8, dk-8000, aarhus c, denmark 2 department of geosciences and natural resource management, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark 3 department of biology, university of copenhagen, øster voldgade 5–7, dk-1350 copenhagen k, denmark. in front of the delta, numerous branches and twigs were observed to be floating on the water. we speculated as to whether they were recent, having been eroded by rivers during the growth of the glacier – or if they were mid-holocene, from a time when the greenland ice sheet was smaller than at present (briner et al. 2016). radiocarbon dating yielded modern ages (not shown) indicating that the plant remains are modern. in the deeper parts of the lake, more than 50 m of sediments have accumulated (fig. 3b). if we assume that sedimentation began c. 10 ka ago, this would correspond to a mean sedimentation rate of 0.5 cm per year, which we consider typical for a proglacial lake. the lake is frozen for most of the year but ice-free in the summer. its large size and location near the margin of the ice sheet, means that strong winds can develop over the water, leading to fairly large waves, which prevent fine-grained sediments from accumulating in shallow areas. the sediments appear acoustically laminated (fig. 2d) with parallel reflections in the seismic profiles and the sediments drape over the underlying irregular bedrock surface (fig. 4). plough marks are visible in some seismic profiles, at depths of up to c. 40 m (fig. 4a). no icebergs occur in the lake at present, which might otherwise have caused the plough marks. however, a local hunter who visited the area in 1862 and 1865 noted the presence of icebergs in the lake, hence the glacier was calving into the lake at that time (barselaj 1866; weidick et al. 2012). the plough marks are not covered by sediments and we consider it likely that they formed during the little ice age. conclusions lake tasersuaq contains more than 50 m of fine-grained glacial rock flour in the deeper parts of the lake. the sediment accumulated following the last deglaciation, around 10 000 years ago. the mapping carried out in 2018 provides the first understanding of the spatial distribution and amount of rock flour deposits in the lake and forms an important first step to assessing the resource potential of these deposits in greenland. further work is now needed to characterise the sediments in terms of their grain size, mineralogy and chemistry and to assess the effectiveness of rock flour as a fertiliser in nutrientpoor soils. acknowledgements geocenter denmark financed the study. lars-georg rödel took care of the technical aspects of the field work. we thank the reviewers, niels tvis and asger ken pedersen, whose comments improved the manuscript. references barselej 1866: taserssualiarnermik kingornalo angalanermik tusagagssiat (account from a voyage to the taserssuaq area). the newspaper atuagagdliutit, 30 november, columns 1238–1240 and 8 december, columns 1246–1251. briner, j.p. et al. 2016: holocene climate change in arctic canada and greenland. quaternary science reviews 147, 340–364. https://doi. org/10.1016/j.quascirev.2016.02.010 escher, j.c. & pulvertaft, t.c.r. 1995: geological map of greenland 1: 2 500 000. copenhagen: geological survey of greenland. larsen, n.k., funder, s., kjær, k.h., kjeldsen, k.k., knudsen, m.f. & linge, h. 2014: rapid early holocene ice retreat in west greenland. quaternary science reviews 92, 310–323. https://doi.org/10.1016/j. quascirev.2013.05.027 ray, d.k., müller, n.d., west, p.c. & foley, j.a. 2013: yield trends are insufficient to double global crop production by 2050. plos one 8, e66428. https://doi.org/10.1371/journal.pone.0066428 van vuuren, d.p., bouwman, a.f. & beusen, a.h.w. 2010: phosphorus demand for the 1970–2100 period: a scenario analysis of resource depletion. global environmental change 20, 428–439. https://doi. org/10.1016/j.gloenvcha.2010.04.004 weidick, a. 1971: quaternary map of greenland, 1:2 500 000. copenhagen: geological survey of greenland. weidick, a. 1978: kvartærgeologisk kort over grønland, frederikshåbs isblink – søndre strømfjord, 1:500 000. copenhagen: geological survey of greenland. weidick, a., bennike, o., citterio, m. & nørgaard-pedersen, n. 2012: neoglacial and historical glacier changes around kangersuneq fjord in southern west greenland. geological survey of denmark and greenland bulletin 27, 68 pp. wise, t.a. 2013: can we feed the world in 2050? a scoping paper to assess the evidence. global development and environment institute working paper 13-04, 36 pp. medford: tufts university. how to cite bennike, o., jensen, j.b, sukstorf, f.n. & rosing, m.t. 2019: mapping glacial rock flour deposits in tasersuaq, southern west greenland. geological survey of denmark and greenland bulletin 43, e2019430206. https://doi.org/10.34194/geusb-201943-02-06 e2019430206-05 mailto:obe%40geus.dk?subject= https://doi.org/10.1016/j.quascirev.2016.02.010 https://doi.org/10.1016/j.quascirev.2016.02.010 https://doi.org/10.1016/j.quascirev.2013.05.027 https://doi.org/10.1016/j.quascirev.2013.05.027 https://doi.org/10.1371/journal.pone.0066428 https://doi.org/10.1016/j.gloenvcha.2010.04.004 https://doi.org/10.1016/j.gloenvcha.2010.04.004 https://doi.org/10.34194/geusb-201943-02-06 geological survey of denmark and greenland bulletin 1, 611-630 611 maturation of organic matter with increasing temperature is one of the most important parameters when evaluating the thermal history and hydrocarbon potential of a basin. a common way to assess the maturation depthtrend is to construct coalification curves by plotting measured huminite/vitrinite reflectances against depth. various attempts have been made to construct such curves for the danish basin and the fennoscandian border zone (thomsen 1980; thomsen et al. 1983, 1987; schmidt 1985). however, these attempts have to some extent overlooked the significant amounts of post-early cretaceous differential uplift that have influenced the danish basin and the fennoscandian border zone. forchhammer (1835) recognised the pronounced erosional unconformity at the base of the quaternary. the hiatus at this surface increases significantly towards the northern and eastern margin of the danish basin. it is now widely accepted that significant regional uplift occurred in neogene and pleistocene times in the north atlantic area and along the norwegian west coast into the burial depth and post-early cretaceous uplift of lower–middle jurassic strata in the fennoscandian border zone based on organic maturity henrik i. petersen, lars h. nielsen,torben bidstrup and erik thomsen the burial depth and the magnitude of late cretaceous – early cenozoic and neogene–pleistocene uplift of lower–middle jurassic strata in the fennoscandian border zone are estimated from measurements of huminite reflectance and comparison with a regional coalification gradient. the regional coalification curve is constructed by plotting uplift-corrected sample depths against more than 300 huminite/vitrinite reflectance values from upper triassic – lower cretaceous deposits in the danish basin and the fennoscandian border zone. the present sample depths are corrected for late cretaceous inversion in the sorgenfrei–tornquist zone and for neogene–pleistocene regional uplift. a coalification curve is erected; it cuts the abscissa at 0.2 %ro corresponding to the reflectance of peat. this curve is considered to approximate to a reliable coalification profile over much of the study area. the jurassic coals from the fennoscandian border zone are of low rank and, based on the regional coalification curve, they have been buried to c. 625–2450 m. in the eastern part of the rønne graben, in the kolobrzeg graben and in the arnager–sose fault block, the jurassic strata were subsequently uplifted c. 290–1400 m, corresponding to the amount of late cretaceous – early cenozoic inversion observed on seismic sections. thus, it appears that neogene–pleistocene uplift did not influence the bornholm area significantly. the data from the höganäs basin and fyledal indicate a total uplift of c. 1450–2450 m, corresponding to estimates from the inversion zone in the kattegat. the data from anholt, on the eastern margin of the inversion zone, indicate c. 975 m of uplift. keywords: danish basin, fennoscandian border zone, lower–middle jurassic, organic maturity, coalification curve, burial depth, late cretaceous – early tertiary inversion, neogene–pleistocene uplift geological survey of denmark and greenland, geocenter copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: hip@geus.dk geological survey of denmark and greenland bulletin 1, 611–630 (2003) © geus, 2003 612 skagerrak and the kattegat to the swedish coast (manum & throndsen 1978; jensen & michelsen 1992; jensen & schmidt 1992, 1993; nyland et al. 1992; japsen 1993, 1998; michelsen & nielsen 1993). in addition to the regional neogene–pleistocene uplift, significant inversion occurred in the sorgenfrei–tornquist zone in late cretaceous – early cenozoic times due to right-lateral transpression (liboriussen et al. 1987; norling & bergström 1987; eugeno-s working group 1988; michelsen & nielsen 1991, 1993; mogensen 1994). the amount and distribution of the two types of uplift of the danish basin and sorgenfrei–tornquist zone have been estimated by comparing the lateral variation of the sonic velocities of a uniform unit of lower jurassic marine mudstones with the structural and stratigraphic development of the area (jensen & michelsen 1992; japsen 1993; michelsen & nielsen 1993). these estimates indicate uplift of 100–2000 m, emphasising that variable amounts of uplift have to be considered when huminite/vitrinite reflectances are used to interpret the thermal history and predict maturity of undrilled sections. we have used 311 reflectance measurements (corrected for uplift) from 20 well-sections (fig. 1) in order to construct a reliable coalification curve which can be used to evaluate the thermal history and maturation of organic matter in the danish basin and the fennoscandian border zone. the amount of uplift was primarily estimated from seismic interval velocities (japsen 1993). the aims of the paper are: (1) to construct a regional coalification curve for the danish basin and the fennoscandian border zone and (2) to use the curve to estimate the burial depth and later uplift of lower–middle jurassic strata exposed or cored in shallow wells in the fennoscandian border zone by comparing huminite reflectances with the curve. the maturation of organic matter provides the best estimate of the burial and uplift history, as measurements of sonic velocities in general are not available from these localities. geological setting the fennoscandian border zone is divided into the skagerrak–kattegat platform and the sorgenfrei–tornquist zone (fig. 1; sorgenfrei & buch 1964; eugeno-s børglum fault ullerslev-1 horsens-1 rønde-1 voldum-1 terne-1 gassum-1 hobro-1 farsø-1 års-1hyllebjerg-1 vedsted-1 vinding-1 mors-1 kvols-1 børglum-1 haldager-1 frederikshavn-1 skagen-2 sæby-1 lavø-1 stenlille-1 hans-1 ramlösa-1b pernille-1 stina-1 fjerritslev-2 anholt-3/4 fyledal 100 km bornholm skåne fjerritslev fault fjerritslev trough well normal fault erosion limit of triassic–jurassic areas of late cretaceous and early tertiary inversion basement high outcrop locality rø nn e g ra be n skagerrak kattegat ringkøbing–fyn high danish basin skagerrak–kattegat platform sorgenfrei–tornquist zone 56°n 57°n 55°n east north sea high 5 157 18øresund wells øresund 13 8°e 10°e 12°e 13°e fig. 1. structural map showing the danish basin and the fennoscandian border zone, and locations of the study wells and the outcrop at fyleverken sand pit, fyledal, skåne (see also fig. 9). inset map shows the location of the wells in the helsingør–helsingborg area. studied wells and outcrops on bornholm are shown in fig. 6. modified from liboriussen et al. (1987) and eugeno-s working group (1988). 613 working group 1988; michelsen & nielsen 1991, 1993). the skagerrak–kattegat platform is a large stable platform area to the north-east where the mesozoic succession onlaps palaeozoic and crystalline basement rocks and thins towards the north-east. the sorgenfrei–tornquist zone is 20–50 km wide and strongly block-faulted. the zone demarcates the stable baltic shield and forms the north-western extension of one of europes most prominent tectonic structures, the tornquist zone. the sorgenfrei–tornquist zone converges at the rønne graben, offshore bornholm, with the teisseyre–tornquist zone which constitutes a tectonic lineament extending to the black sea (ziegler 1982). the rønne graben is a pull-apart basin that formed by dextral wrench-faulting in late carboniferous – early permian times (vejbæk 1985; liboriussen et al. 1987). during the late palaeozoic – mesozoic breakup of the supercontinent pangaea, the tornquist zone was dominated by transtensional stress resulting in subsidence and tilting of fault blocks, and deposition of mesozoic sedimentary successions up to 8 km thick. the sorgenfrei–tornquist zone demarcates the transition from the baltic shield to the wnw–ese-trending intracratonic danish basin that was formed during a late carboniferous – early permian rift phase accompanied by formation of both extrusive and intrusive volcanic rocks. after rifting, the basin subsided due to thermal cooling (sørensen 1986; vejbæk 1989, 1990). the postrift basin-fill consists of a relatively complete section of zechstein, mesozoic and cenozoic deposits which are 6–7 km thick along the axis of the basin. a gradual shallowing of the basin towards the high-lying basement blocks of the ringkøbing–fyn high to the south-west is indicated by thinning of the section. this is most pronounced for the upper permian (zechstein) strata. the ringkøbing–fyn high was periodically subjected to erosion. in the fjerritslev trough, within the sorgenfrei–tornquist zone, an important thickness anomaly was formed due to transtensional strike-slip movements possibly accompanied by salt withdrawal causing relatively fast subsidence and deposition (bertelsen 1980; liboriussen et al. 1987; vejbæk 1990; christensen & korstgård 1994). late early jurassic – early middle jurassic uplift of the ringkøbing–fyn high, presumably connected to the updoming of the central north sea, caused north-eastwards tilting of the basin and erosion both on the high and in the southern part of the basin (michelsen 1978; koch 1983; nielsen 1993, 1995, 2003, this volume). extensive volcanism and uplift occurred in skåne, possibly connected to general updoming (klingspor 1976; norling & bergström 1987; f. surlyk, personal communication 1996). in late middle u m l u l u l u m l u l u l annero fm a nn er o fm m ar ie da l f m vilhelmsfält fm fyledal clay fortuna marl glass sand mb fuglunda mb r ya f m h ög an äs fm helsingborg mb döshult mb pankarp mb katslösa mb rydebäck mb röddinge fm ? sorthat fm bagå fm hasle fm galgeløkke mb sose bugt mb munkerup mb r øn ne f m gassum fm fjerritslev fm haldager sand fm hettangian sinemurian pliensbachian toarcian aalenian bajocian bathonian callovian oxfordian m id dl e u pp er lo w er ju ra ss ic system stage sw ne nw se danish basin bornholm onshore skåne flyvbjerg fm ? ?? ? fig. 2. stratigraphic scheme of the lower–middle jurassic of the danish basin, bornholm and skåne. compiled from nielsen (1995; 2003, this volume), ahlberg et al. (2003, this volume) and michelsen et al. (2003, this volume). – late jurassic times, the ringkøbing–fyn high began to subside again and gradually the basin expanded towards the south-west, attaining its previous size towards the end of the early cretaceous (nielsen 1995). during late cretaceous – early cenozoic times, transpressional tectonism in the sorgenfrei–tornquist zone and the rønne graben, caused by the alpine orogenesis, led to reactivation and pronounced inversion of fault blocks in the zone, which together with neogene– pleistocene regional uplift resulted in erosion of the mesozoic deposits (gry 1969; gravesen et al. 1982; liboriussen et al. 1987; norling & bergström 1987; japsen 1993, 1998; michelsen & nielsen 1991, 1993). the lower–middle jurassic stratigraphy of the danish basin, bornholm and skåne is shown in figure 2. note that the upper pliensbachian – lower aalenian section on bornholm has previously been referred to the lower bagå formation (koppelhus & nielsen 1994), but is now assigned to the sorthat formation (michelsen et al. 2003, this volume). prerequisites for construction of a regional coalification curve construction of a regional coalification curve requires that the following conditions are fulfilled: (1) many geographically widespread data, (2) exclusion of data from well-sections with an abnormal thermal history, (3) reliable corrections for uplift and (4) a relatively constant temperature gradient both in time and space. although totalling more than 300, the reflectance measurements from the triassic – lower cretaceous deposits in the danish basin and the fennoscandian border zone only partly fulfil the first condition (fig. 1). close to half of the data points come from the central part of the basin from wells such as mors-1, hyllebjerg-1, års-1 and farsø-1. however, this is not a serious problem because in this area there is good agreement between uplift values determined from shale velocities and results from modelling of the maturation history. for example, in the hyllebjerg-1 well (fig. 3), the uplift determined by basin modelling is 600 m compared with 575 m obtained from shale velocities. reflectance values from deposits directly overlying salt diapirs clearly indicate a locally increased heat flow (uglev-1; schmidt 1985), and have been omitted in the construction of the regional curve. the corrections for uplift of the selected well-sections have been mainly calculated from seismic velocity data (japsen 1993). the amount of correction is further controlled by the well-known stratigraphy and structures of the area as shown by interpreted seismic sections and preserved thicknesses of lithostratigraphic units in wellsections. the basin subsided due to thermal contraction after the permian rift phase with raised heat flow (vejbæk 1989). the initial phase of the basin development was probably characterised by lateral variations in heat flow, but after the deposition of a thick sedimentary cover of zechstein and lower–middle triassic deposits, the lateral variations in heat flow probably ceased due to the effect of sediment blanketing (nielsen & balling 1990). recent mapping and modelling of the subsurface temperature at various depths demonstrate a uniform regional temperature distribution in the study area (j.j. møller, personal communication 1997). furthermore, regional corrected measurements of heat flow at the sur614 0 1600 3200 d ep th ( m ) 1.200.600.00 %r : measured %ro : modelled %r fig. 3. modelled reflectance values and selected measured reflectance values from the hyllebjerg-1 well. the uplift determined by basin modelling is 600 m compared to 575 m obtained from shale velocities. this confirms that the area around hyllebjerg-1 is a good reference area. face are uniform, in the range 60–70 mw/m2 (balling 1995). based on the structural and thermal evolution of the basin, therefore, we assume that the lateral variation in heat flow from late triassic to recent times in general was relatively small, and that lateral heat flow variations did not influence the maturation of the organic matter significantly. the necessary prerequisites for the construction of a regional coalification curve are thus fulfilled to a large extent. in order to avoid complications due to the late early jurassic – middle jurassic uplift of the ringkøbing–fyn high and the southern parts of the basin, the majority of the data used for construction of the curve comes from well-sections that experienced continued subsidence or only limited uplift and erosion during this phase (nielsen 1995). influence of temperature on maturation the increase in rank with depth in a well-section is mainly caused by rising temperature with depth, and the rate of rank increase is strongly dependent on the geothermal gradient. in a sandstone succession, for instance, the rank gradient is much lower than in a mudstone succession due to the higher thermal conductivity of sandstones (damberger 1968). raised heat flow and thus a higher geothermal gradient also raises the rank gradient (teichmüller 1979; suggate 1998). however, in a basin with a relatively constant heat flow and a laterally uniform lithology, the rank of coals is primarily dependent on the maximum temperature to which the organic matter was subjected, which corresponds to the maximum burial depth. since coalification of organic matter is irreversible, later uplift does not influence the measured reflectance values. as demonstrated below, the majority of the reflectance values from the danish basin and the fennoscandian border zone are interpreted to show a normal coalification trend. the rank of the coals thus reflects the maximum burial depth of the strata. construction of the coalification curve a total of 311 reflectance measurements from the upper triassic, jurassic and lower cretaceous successions in twenty wells from the fennoscandian border zone and the danish basin are used in the construction of the coalification curve (fig. 1). all the reflectance values are random measurements. samples from drill cores were available from the frederikshavn-1, fjerritslev-2, gassum-1, haldager-1, horsens-1, lavø-1, skagen-2, ullerslev-1, vedsted-1 and vinding-1 wells. where possible, samples with coal, coaly inclusions, or dark-coloured shales or siltstones were selected for reflectance measurements as these rocks contain the best organic matter for rank determination (thomsen 1980). untreated rock samples were studied as they more easily allow identification of oxidised and bituminous organic matter (unsuitable for reflectance measurements) than is the case for kerogen concentrates. in coal or coaly samples, approximately 50–100 measurements were carried out in each sample; in samples with disseminated organic matter, as many particles as possible were measured, usually 20–50. core samples from the børglum-1 well were analysed in the same fashion (schmidt 1985), whereas coaly or dark-coloured shaly intervals were selected from sidewall cores and core samples in the farsø-1 well (thomsen 1983). cuttings and a limited number of core samples were available from the års-1 well, whereas only cuttings were available from the rest 615 års-1 børglum-1 farsø-1 fjerritslev-2 frederikshavn-1 gassum-1 haldager-1 hans-1 hobro-1 horsens-1 hyllebjerg-1 kvols-1 lavø-1 mors-1 rønde-1 skagen-2 terne-1 ullerslev-1 vedsted-1 vinding-1 558 m 1273 m 481 m 1531 m 1000 m 1190 m 1400 m 1733 m 550 m 200 m 575 m 448 m 1000 m 785 m 513 m 1100 m 1405 m 500 m 1400 m 250 m total post-early cretaceous uplift (amount of correction) comments on the estimated amounts of uplift well table 1. correction for post-early cretaceous uplift after japsen (1993) after japsen (1993) after japsen (1993) after japsen (1993) by comparison to sæby-1, 997 m (japsen 1993) by comparison to voldum-1, 942 m (japsen 1993) and 248 m deeper truncation of chalk by comparison to børglum-1 and fjerritslev-2 after japsen (1993) by comparison to års-1 by using fig. 1c in japsen (1993) after japsen (1993) after japsen (1993) by using fig. 1c in japsen (1993) and stenlille-1, 850 m after japsen (1993) after japsen (1993) by comparison to sæby-1, 997 m after japsen (1993) by using fig. 1c in japsen (1993) by comparison to børglum-1 and fjerritslev-2 by comparison to mejrup-1, 327 m, and vemb-1, 309 m (japsen 1993) of the wells (rønde-1: thomsen 1980; hyllebjerg-1: schmidt 1988; hans-1, hobro-1, kvols-1, mors-1, terne-1: geus, unpublished data). all sample depths were corrected for post-early cretaceous net uplift before being plotted against reflectance (table 1). the correction for the well-sections was based on the analysis of sonic velocities of shales by japsen (1993). the main uncertainties related to this method are the requirement 616 0.0 0.2 0.4 0.6 0.8 0 500 1000 1500 2000 2500 3000 3500 4000 %ro d ep th ( m ) års-1 børglum-1 farsø-1 fjerritslev-2 frederikshavn-1 gassum-1 haldager-1 hans-1 hobro-1 horsens-1 hyllebjerg-1 kvols-1 lavø-1 mors-1 rønde-1 skagen-2 terne-1 ullerslev-1 vedsted-1 vinding-1 0.0 0.2 0.4 0.6 0.8 %ro års-1 farsø-1 frederikshavn-1 gassum-1 hans-1 hobro-1 horsens-1 hyllebjerg-1 kvols-1 lavø-1 mors-1 rønde-1 skagen-2 ullerslev-1 vinding-1 a b fig. 4. a: coalification profile for the 20 investigated wells (n = 311); the depths are corrected for post-cretaceous uplift. data from thomsen (1980, 1983, 1984), schmidt (1985, 1988) and unpublished geus data. b: coalification profile (uplift-corrected depths) for 15 wells in the danish basin, the fennoscandian border zone and the skagerrak–kattegat platform (n = 249). the regression line has a correlation coefficient of 0.88, and intercepts the reflectance axis at 0.21 %ro. of a uniform shale unit covering the entire study area and a valid reference curve (discussed in japsen 1993). the analyses of the velocity data provide an uplift pattern that is compatible with structural interpretations of seismic sections and stratigraphic well data. the present sample depths are thus corrected by adding the amount of uplift proposed by japsen (1993). the uplift of wells not included by japsen (1993) is estimated by comparison to nearby wells and interpolation (table 1). the gassum-1 well, for example, is situated on a salt structure, in a setting comparable to the nearby voldum-1 well, which was uplifted 942 m (japsen 1993). the chalk section in gassum-1 seems to be more deeply truncated than that in voldum-1 (c. 250 m), and an uplift of c. 1190 m is thus estimated for the gassum-1 section. the uplift of the horsens-1 and ullerslev-1 wells is estimated to be 200 m and 500 m, respectively, by using the general uplift map of japsen (1993). the uplift of the lavø-1 well is estimated to be 1000 m based on the same map and comparison with the stenlille-1 well (850 m). two coalification trends are recognised when the reflectance measurements are plotted against uplift-corrected depths: a well-defined main trend to the right and a less pronounced trend to the left (fig. 4a). the main coalification trend is based on well data from the danish basin and the skagerrak–kattegat platform (fig. 4b), whereas well data from the fjerritslev trough (børglum-1, fjerritslev-2, haldager-1, vedsted-1) and the terne-1 well define the other coalification trend. huminite/vitrinite reflectance is largely unaffected by the formation of mainly isolated aromatic rings up to a reflectance of c. 0.7 %ro, resulting in a slow and more or less linear increase in reflectance (carr & williamson 1990; suggate 1998). however, as maturation proceeds past this point the formation of polycyclic aromatic units significantly increases the reflectance. as the level of coalification in the studied sections is in the low rank range (< 0.7 %ro), it is possible to approximate this ‘straight’ section of the curve with a linear regression line. the linear regression line for the danish basin/skagerrak–kattegat platform intersects the %ro-axis at 0.2 %ro, which is the expected reflectance at the surface (fig. 4b; dow 1977; suggate 1998). the rank gradient is 0.12 %ro/km. the absence of significant anomalies indicates that the assumption of a relatively uniform geothermal gradient in the study area is justified, and that the corrections for uplift are reasonable. the regression line of the coalification curve is based on 249 samples; it has a correlation coefficient of 0.88, and is considered to be a good approximation of the coalification profile. this relationship can then be used to estimate the maximum burial depths and later uplift of jurassic coal seams in the fennoscandian border zone. the reasonable and stratigraphically consistent estimates are taken as indirect evidence of the reliability of the coalification curve, and the coalification curve is considered as a valid approximation of the maturation trend over most of the study area. the data from børglum-1, fjerritslev-2, haldager-1, vedsted-1 and terne-1 form an atypically steep coalification trend which was investigated by modelling of the basin development. modelled vitrinite reflectances for the børglum-1 well are shown in figure 5, together with the measured values. the basin modelling indicates uplift of only 800 m, compared with 1300 m derived from shale velocities. interval velocities for the 617 0 1600 3200 d ep th ( m ) 1.200.600.00 %r : measured %ro : modelled %r fig. 5. modelled reflectance values and measured reflectance values from the børglum-1 well. the amount of uplift determined by basin modelling is 800 m compared to the 1300 m obtained from shale velocities. it is suggested that the uplift of the wells in the fjerritslev trough (and the terne-1 well) determined from the shale velocities is overestimated due to the sand-rich nature of the succession in these wells. chalk section in the fjerritslev trough are low, and new data based on chalk group velocities yield significantly lower uplift values (japsen 1998). the data suggest correction for uplift of only 817 m for børglum-1, 1012 m for fjerritslev-2, and 644 m for haldager-1. if the present depths of these wells are uplift-corrected according to the new chalk group data, the huminite reflectance values would fit the coalification curve shown in figure 4b. this suggests that uplift values determined from shale velocities are probably overestimates. the relatively high shale velocities in this area (japsen 1993) may be the result of an increased coarse-grained component within the lowermost fjerritslev formation than is typical of the rest of the basin. estimation of burial depth and later uplift of lower–middle jurassic coal seams in the fennoscandian border zone determination of coal rank the rank, or the level of thermal maturity, of the coals was determined by random reflectance measurements on eu-ulminite, the brown coal equivalent of collotelinite, following the standard procedure outlined in taylor et al. (1998). up to 100 measurements were made on each sample and a mean reflectance (%ro) was calculated. parameters obtained by organic geochemical methods, when available, were used to support the estimation of the maturity. rock-eval pyrolysis was used to determine the tmax value which is the temperature at which maximum pyrolysis occurs (peters 1986). the value of tmax increases with increasing maturity of the organic matter, but is also influenced by the composition of the organic matter, as a high content of inert components may result in an elevated tmax value. biomarker data were obtained from the saturated fraction of solvent extracts from the coal samples using gas chromatography/mass spectrometry. pentacyclic hopanes were identified by detecting the most characteristic fragmentation ion, the m/z 191. the occurrence of the thermally unstable hopenes and 17ß(h),21ß(h) forms in the coals is indicative of immaturity, and the c31-homohopane 22s/(22r+22s) epimerisation ratio approaches an equilibrium value of approximately 0.6 with increasing maturity. the sensitivity of the hopanes to thermal influence at low maturity levels may be a valuable criterion when evaluating the validity of maturity differences obtained by reflectance measurements. the carbon preference index (cpi) was calculated from traces obtained from gas chromatography. the cpi is a measure of the predominance of odd carbon numbered n-alkanes calculated over a specified range (bray & evans 1961; cooper & bray 1963). in the present study, the range was from nc22 to nc32. the dominance of odd carbon numbered n-alkanes together with a prominent heavy-end fraction are characteristic features of terrestrially derived organic matter (isaksen 1995). however, increasing maturity seems to minimise the cpi due to alkane cleavage reactions and dilution of the odd carbon numbered n-alkanes (radke et al. 1980). hence, the cpi tends to reach an equilibrium of 1 with increasing maturity. huminite reflectances of the coals the following section focuses on the degree of coalification of lower and middle jurassic coal seams and strata from the island of bornholm, the rønne and kolobrzeg grabens (offshore bornholm), skåne, the øresund area, and kattegat in order to present the rank distribution. hettangian–sinemurian rønne formation, bornholm two coal seams in the hettangian munkerup member of the rønne formation in the arnager–sose fault block have reflectances of 0.40 %ro and 0.41 %ro (figs 2, 6; table 2). however, samples from carbonaceous seams in the overlying hettangian–sinemurian sose bugt member display significantly lower reflectance values (0.28 %ro and 0.30 %ro) (figs 2, 6; table 2). the coal seam in the sinemurian galgeløkke member of the rønne formation at the galgeløkke coastal cliff (rønne–hasle fault block) has a reflectance of 0.38 %ro (figs 2, 6; table 2). the galgeløkke-2 well was drilled close to the galgeløkke cliff and cored c. 400 m of hettangian–sinemurian strata (nielsen 1995). a hettangian coal seam situated at a depth of 383 m has a reflectance value of 0.42 %ro. this is compatible with a tmax value of 427°c and a c31-homohopane 22s/(22s+22r) epimerisation ratio of only 0.06. the coal seam is estimated to lie 50–75 m above the base of the jurassic, as suggested by log correlations between the galgeløkke-2, pernille-1 and stina-1 wells (nielsen 1995). 618 619 rønne– hasle fault block nyker block rønne graben arnager–sose fault block gudhjem trough bornholm high 5 6 3 4 2 1 10 km c t n fig. 6. geological map of bornholm showing the investigated localities: 1, exposed munkerup member (rønne formation); 2, type section of the sose bugt member (rønne formation); 3, type section of the galgeløkke member (rønne formation) at the galgeløkke coastal exposures and the galgeløkke-2 well; 4, korsodde coastal cliff, sorthat formation; 5, levka-1 well, sorthat formation; 6, hasle klinkerfabrik clay pit, upper bagå formation. modified from gravesen et al. (1982). sose bugt section hettangian–sinemurian 0.29 2 625 anholt well* lower jurassic 0.33 12 975 975 levka-1 well / upper pliensbachian – 0.36 13 420(60) 0.06–0.1 1210 1210 korsodde section** lower toarcian øresund-15 well+ sinemurian 0.36 1 419(1) 1210 1210 galgeløkke section sinemurian 0.38 2 1400 1400 øresund-5 well+ bajocian 0.39 1 421(2) 1450 1450 fyleverken sand pit bajocian 0.39 2 430(2) 1450 1450 munkerup section hettangian 0.41 2 1660 1660 galgeløkke-2 well hettangian 0.42 1 427(1) 0.06 1700 1320 hasle kl. clay pit bathonian 0.42 6 1700 1700 øresund-7 well+ bajocian–bathonian 0.43 5 418(9) 0.05–0.09 1875 1875 øresund-13 well+ hettangian–sinemurian 0.46 1 426(1) 0.23 2100 2010 øresund-18 well+ hettangian–sinemurian 0.51 2 426(7) 0.37–0.42 2440 2440 # the average reflectance is calculated from mean random reflectance (nos of analyses). a mean random reflectance value is based on c. 100 measurements in each sample. * data from nielsen et al. (2003, this volume). ** data from petersen & nielsen (1995). + data from petersen (1994). ++ post-early cretaceous uplift: 290 m (petersen et al. 1996) locallity table 2. maturity data, estimated burial depth/magnitude of later uplift age average reflectance %ro # average tmax (c°) (n) net uplift (m) maximum burial depth (m) c31-homohopane 22s/(22s+22r) number of mean random reflectances 625++ cretaceous lower–middle jurassic sorthat formation and middle jurassic bagå formation lower jurassic hasle formation lower jurassic rønne formation upper triassic kågeröd formation lower palaeozoic precambrian crystalline basement fault 620 upper pliensbachian – lower toarcian part of the sorthat formation, bornholm reflectance values of coal seams from the upper pliensbachian – lower toarcian part of the sorthat formation in cores from the levka-1 well and in the korsodde coastal cliff (rønne–hasle fault block) average 0.36 %ro (figs 2, 6; table 2). tmax values average 421°c, and c31-homohopane 22s/(22s+22r) epimerisation ratios for eight samples from the levka-1 well are in the range 0.06–0.1, which conforms with low rank coals. thermally sensitive hopenes and ßß-forms occur, and cpi values greater than 3 are compatible with immature low rank coals (petersen & nielsen 1995). bathonian part of the bagå formation, bornholm the coal-bearing aalenian–bathonian bagå formation is exposed in the hasle klinkerfabrik clay pit (figs 2, 6). the clay pit is situated in the rønne–hasle fault block, close to the faulted margin of the rønne graben. reflectance values from eight coal seams in the uppermost part of the bagå formation (bathonian) are in the range 0.36–0.46 %ro, averaging 0.42 %ro (table 2). lower jurassic in the rønne and kolobrzeg grabens the lower jurassic from the western part of the rønne graben is represented by the pernille-1 well, and from the kolobrzeg graben by the stina-1 well (fig. 7). in pernille-1, the lower jurassic from 980–1470 m is referred to the hettangian–sinemurian rønne formation (nielsen 1995). the reflectance values show a well-defined coalification trend with increasing depth (fig. 8). in stina-1, the lower jurassic from 180–560 m possibly includes the hettangian to upper pliensbachian (nielsen 1995). the reflectance values vary between 0.31 %ro and 0.40 %ro (fig. 8). hanö bay half-graben christiansø ridge ustka block darlowo block kolobrzeg graben gryfice graben rønne graben colonus trough romele ridge vomb trough skurup block bornholm gat block ar na ge r– so se bl oc k ri se bæ k tr ou gh arkona block w iek–trent block wolin block bornholm high gudhjem trough rø nn e fa ult g at fa ult trzebiatow fault adler fault koszalin–chojnice fault ustki fault h am m er fault ringsjö–andrarum fault fyledal fault hälsingborg–romele fault 25 km 55°n 14°e 15°e 16°e pernille-1 stina-1 fig. 7. structural map showing the block mosaic of the bornholm region and the position of the pernille-1 and stina-1 wells. the relative significance of the faults is indicated by varying line thickness. slightly modified from hamann (1994). 621 bajocian fuglunda member of skåne bajocian coal-bearing sediments of the fuglunda member (mariedal formation) are exposed in the fyleverken sand pit, fyledal at eriksdal (figs 1, 2). the sediments are part of a 400–600 m thick succession that was deposited close to the boundary between the vomb trough and the palaeozoic colonus trough (fig. 9). two coal samples from the lower part of the fuglunda member show reflectance values of 0.38 %ro and 0.39 %ro (table 2), and tmax values of 427°c and 433°c. jurassic of the northern øresund knowledge of the jurassic in the helsingør–helsingborg area is derived from sea cliff and quarry exposures, from tunnel excavations, and from cored shallow wells in the øresund and skåne (fig. 1). the deposits belong to the faultand flexure-bounded höganäs basin within the sorgenfrei–tornquist zone (fig. 9). hettangian–sinemurian sediments occur in the øresund-13 and -18 wells. a coal seam in the former well yielded a reflectance of 0.46 %ro, whereas a coal seam in the latter well yielded a reflectance of 0.51 %ro (table 2). a sinemurian coal seam in the øresund-15 well has a reflectance of 0.36 %ro (table 2). the relatively low reflectance value of the seam from the øresund-15 well may be suppressed, as the pronounced content of huminite and the presence of pyrite indicate 0 500 1000 1500 2000 d ep th ( m ) 0.60.4 %ro 0.20.0 pernille-1 stina-1 stina-1 corrected vomb trough colonus trough n fyledal fault fyleverken sand pit rom ele ridge ▲ ▲ ■■ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ■■ ■ ■ ▲ ▲ ■■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ■■ ■■ ■ ■ ■■ ■ ■ ■ ■ ■■ ■■ ■■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ 25 km höganäs basin skurup block rhaetian and jurassic present distribution reverse fault normal fault fault of unknown type fig. 8. coalification profiles for the lower jurassic in the pernille-1 well in the rønne graben and the stina-1 well in the kolobrzeg graben. the regional coalification profile for the danish basin/ skagerrak–kattegat platform is also shown. the profile from pernille-1 falls on the regional trend indicating that the section has not been uplifted. the stina-1 section has been uplifted c. 900 m. fig. 9. structural map of skåne showing the present distribution of rhaetian and jurassic deposits. modified from norling & bergström (1987). that anoxic and possibly occasional saline conditions were present in the precursor mire. this may have resulted in the formation of hydrogen-enriched huminite, which matures at a lower rate (petersen & rosenberg 1998). the tmax is also low (419°c). the tmax value of the coal from the øresund-13 well is 426°c and the values from the øresund-18 well are 424–428°c. the c31-homohopane 22s/(22s+22r) epimerisation ratio is 0.23 for the øresund-13 well, whereas the ratio is in the range 0.37–0.42 for the øresund-18 well. together with the reflectance values, these data are consistent with a higher rank of the coals from the øresund-18 well compared to the coal from the øresund-13 well. the bajocian–bathonian coal from the øresund-5 well has a reflectance of 0.39 %ro, and the coals from the øresund-7 well have reflectances in the range 0.41–0.46 %ro, averaging 0.43 %ro (table 2). relative proportions of the thermally unstable hopenes and ßßforms are high in the øresund-7 coals, and the c31homohopane 22s/(22s+22r) epimerisation ratios are low (0.05–0.09). the average tmax value and the hopane ratios are similar to the values obtained from the upper pliensbachian – lower toarcian (sorthat formation) coal seams in the levka-1 well on bornholm. however, the reflectance values of the coal seams from the øresund-7 well are higher, possibly due to increased oxidation during the humification process in the early stages of biochemical gelification (diessel 1992; hao & chen 1992). lower–middle jurassic fjerritslev and haldager sand formations, kattegat approximately 200 m of lower and middle jurassic deposits of the fjerritslev and haldager sand formations overlain by c. 100 m of quaternary deposits were cored by a shallow well on the island of anholt located in the kattegat (figs 1, 2). reflectance measurements by nielsen et al. (2003, this volume) give an average value of 0.33 %ro (table 2). estimation of burial depths and later uplift the average reflectance values of the coals and the constructed regional coalification curve are used to estimate the maximum burial depth of the coal-bearing lower and middle jurassic strata in the fennoscandian border zone. the amount of later uplift equals the maximum burial depth in most cases as the majority of the samples were collected close to present sea level. bornholm a burial depth of c. 1660 m is suggested for the outcropping coals of the hettangian munkerup member, whereas the data from the overlying hettangian to sinemurian sose bugt member (both of the rønne formation) indicate only c. 625 m of burial (fig. 10; table 2). the thickness of the non-exposed section between the seams in the munkerup member and those in the sose bugt member is less than a few tens of metres (gry 1969; gravesen et al. 1982); a difference of approximately 1000 m in estimated burial depth between the seams is therefore impossible. the reflectance values from the munkerup member are considered to be too high, possibly due to hydrothermal influence; this aspect is discussed further below. the sose bugt member in the arnager–sose fault block is overlain by at least 60 m of sediments assigned to the hasle formation (gravesen et al. 1982; surlyk et al. 1995), which in turn are unconformably overlain by lower cretaceous sediments. the base of the lower cretaceous was buried to c. 290 m followed by a similar amount of uplift (petersen et al. 1996). this suggests that part of the uplift experienced by the sose bugt member was related to the middle jurassic uplift event, which influenced the areas outside the sorgenfrei–tornquist zone (andsbjerg et al. 2001; nielsen 2003, this volume), and only c. 290 m of the total uplift was caused by post-early cretaceous uplift. the coal seams of the sinemurian galgeløkke member exposed at the galgeløkke coastal cliff have been buried to c. 1400 m with a later uplift of the same magnitude (fig. 10; table 2). the hettangian coal seam at a depth of 383 m in the nearby galgeløkke-2 well was buried to c. 1700 m, corresponding to a burial depth for the base of the jurassic of c. 1750–1775 m with a later uplift of c. 1320 m. the difference in estimated burial depths of the two seams is thus c. 300 m which corresponds reasonably well with the estimated thickness of c. 400 m of the intervening strata. the coal seams in the upper pliensbachian – lower toarcian part of the sorthat formation have been buried to c. 1210 m (fig. 10; table 2). the difference between the estimated burial depths of the seam from the galgeløkke coastal cliff and the seams from the levka-1 and korsodde sections is c. 200 m which conforms well with an estimated thickness of 225–290 m of intervening strata. however, it should be noted that the estimated burial depth of the levka-1 and korsodde coals may be an under-estimate as the reflectance values of the coals may be slightly suppressed (petersen et al. 2003, this volume). 622 623 the data from the bathonian coal seams in the uppermost bagå formation indicate a burial depth of c. 1700 m (fig. 10; table 2). this corresponds to c. 800 m deeper burial than the levka-1 and korsodde sections, and is unrealistic considering the structural and stratigraphic development of the two localities (gry 1969; gravesen et al. 1982; jensen & hamann 1989). hence, it is likely that the high reflectance values of the coals in the upper part of the bagå formation were caused by an additional factor linked to the proximity of the major rønne–hasle fault as discussed below. these estimates suggest that the base of the jurassic in the rønne–hasle block at galgeløkke-2 was buried to c. 1750–1775 m and later uplifted c. 1320 m. this is consistent with the seismic interpretation by hamann (1994) indicating up to 1400 m of inversion along the rønne–hasle fault. the maximum burial depth of 1750–1775 m is compatible with the likely cumulative thicknesses of the jurassic–cretaceous section in the rønne–hasle block. in pernille-1, for instance, the jurassic–cretaceous section has a minimum thickness of c. 1450 m, comprising c. 600 m of lower jurassic deposits overlain by c. 850 m of probable lower to upper cretaceous deposits. in the rønne–hasle block, the cumulative thickness may have been up to 2000 m, as c. 1000 m of jurassic deposits are preserved, and the original thickness of the cretaceous succession may have been up to c. 1000 m, as suggested by reference to the adjacent nyker block and pernille-1. rønne and kolobrzeg grabens the reflectance values of the rønne formation in pernille-1 plot on the trend of the coalification curve (fig. 8). thus the present depth of the jurassic deposits corresponds to the maximum burial depth. this contrasts with the interpretation of a seismic section through the pernille-1 well indicating c. 270 msec (c. 400 m) of inversion relative to the stable skurup block (fig. 11). the data from the stina-1 section situated in the kolobrzeg graben suggest a burial depth of c. 1500 m for the base of the jurassic and c. 900 m of later uplift (figs 7, 8). seismic interpretation suggests a similar sose bugt section (0.29 %ro; hettangian–sinemurian) anholt well (0.33 %ro; lower jurassic) levka-1 well and korsodde section (0.36 %ro; u. pliensbachian – l. toarcian)/ øresund-5 well and fyleverken sand pit (0.39 %ro; bajocian) galgeløkke section (0.38 %ro; sinemurian) munkerup section (0.41 %ro; hettangian) galgeløkke-2 well (0.42 %ro; hettangian)/ øresund-7 well (0.43 %ro; bajocian–bathonian) øresund-13 well (0.46 %ro; hettangian–sinemurian) øresund-18 well (0.51 %ro; hettangian–sinemurian) hasle klinkerfabrik clay pit (0.42 %ro; bathonian) øresund-15 well (0.36 %ro; sinemurian) 0.80.60.40.20.0 %ro 0 500 1000 1500 2000 2500 3000 d ep th ( m ) fig. 10. the standard coalification curve for the danish basin/skagerrak–kattegat platform used to estimate the burial depth of lower and middle jurassic coalbearing strata from the islands of anholt and bornholm, the øresund area and skåne by means of huminite reflectance values. 624 amount of uplift during the inversion phase (hamann 1994; vejbæk et al. 1994). stina-1 is located close to the section shown in figure 12b, which indicates inversion of 1000–1100 m within the kolobrzeg graben. evidence of inversion is also prominent on a sw–ne section east of bornholm (fig. 12c). the lower jurassic strata in the pernille-1 and stina-1 wells were buried to the same depth in the two areas prior to uplift at stina-1. the lower jurassic is unconformably overlain by quaternary deposits in stina-1, whereas the lower jurassic is overlain by c. 850 m of upper cretaceous deposits in pernille-1; this is consistent with the estimate of c. 900 m of uplift at stina-1. fyleverken sand pit, fyledal, skåne the reflectance values of the bajocian coals from fyledal suggest a burial depth of c. 1450 m in the eriksdal area and a similar amount of later uplift (fig. 10; table 2). in the vomb trough, crystalline basement is overlain by up to 200 m of jurassic – lower cretaceous strata and c. 1000 m of upper cretaceous deposits, but much of the santonian–campanian succession accumulated contemporaneously with the inversion of the sorgenfrei– tornquist zone (erlström & guy-ohlson 1994). the thickness of the santonian–campanian succession thus reflects the relative fault movements and not directly the uplift. in the fyleverken sand pit section in fyledal, the post-bajocian jurassic section is c. 400 m thick, and the lower cretaceous – turonian section was probably c. 200 m thick before tilting during santonian–campanian inversion (norling & bergström 1987). erlström (1994) suggested that c. 300 m of maastrichtian deposits were eroded during neogene uplift. the jurassic–cretaceous succession may thus have been c. 900 m thick. paleocene and eocene deposits seem to be confined to the basin south-west of the helsingborg–romele fault. however, cenozoic deposits may have covered eastern skåne and are thought to be present in the hanö bay area (lidmar-bergström 1982; norling & bergström 1987). it is thus very difficult to evaluate in detail the thickness of strata that has been removed on the basis of the preserved stratigraphy, and to support the burial estimate of c. 1450 m. in addition, hydrothermal influence on the rank cannot be precluded (see below). the amount of inversion in the north-western part of the sorgenfrei– tornquist zone is typically suggested to 0 1 2 3 4 t w t ( se c) pernille-1w e basement basement triassic triassic cretaceous jurassic permian l. palaeozoic fig. 11. interpreted seismic section through the pernille-1 well showing c. 270 msec of inversion corresponding to c. 400 m (interval velocity = 2942 m/sec; nielsen & japsen 1992). slightly modified from vejbæk et al. (1994). 625 be up to 1000 m (jensen & michelsen 1992; japsen 1993; michelsen & nielsen 1993). extension of the regional trend of neogene uplift shown by jensen & michelsen (1992) and japsen (1993) into skåne may suggest 1000–1500 m of uplift in this area. however, this trend is mainly based on data to the north-west of skåne and thus may not be directly applicable. northern øresund in the höganäs basin, the data from the hettangian– sinemurian coal seams in the øresund-13 and -18 wells indicate burial depths of c. 2100 m and 2440 m respectively (figs 1, 9, 10; table 2). uplift in the øresund area was evaluated by modelling of a pseudo-well, based on a stratigraphic reconstruction from the shallow wells (fig. 13) and the stratigraphic information given in plate 21 of larsen et al. (1968). the modelling shows that the differences in reflectance (%ro) can be explained by the thickness of the intervening strata with the exception of the øresund-15 well. the low value obtained from this well may be due to suppression, as mentioned earlier. the basin modelling indicates an uplift of 1200 m for the øresund-1 well location, corresponding to uplift of 2100 m at the øresund-13 well location and 2550 m at the øresund-18 well location. this is in close agreement with the values determined from the coalification curve and is compatible with the stratigraphy (fig. 13; larsen et al. 1968, plate 21). reflectance values of three hettangian coal samples from the nearby helsingborg railway tunnel range from 0.48 %ro to 0.57 %ro, and dispersed huminite from hettangian deposits in the nearby ramlösa-1b well gives 0 1 2 3 4 5 0 1 2 3 4 5 skurup block rønne graben arnager block risebæk graben darlowo block hanö bay basin baltic syneclise christiansø high rønne graben ko sz al in f au lt k os za lin f au lt kolobrzeg graben kolobrzeg graben darlowo block ustka block bornholm block svaneke trough christiansø high se ne u. cretaceous jurassic–triassic permian pre-permian fault basement t w t (sec) t w t (sec) t w t (sec) a c c b a nw sw sw ne 20 km b 0 1 2 3 4 5 6 ■■ ■■ ■ ■ ■■ ■ ■ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ■■ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ■■ ■ ■ ■ ■ ▲ ▲ ▲ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■■ fig. 12. geosections around bornholm (see inset map) with depths in two-way travel time (twt), based on interpreted seismic sections. a: note the lack of significant inversion relative to the skurup platform at the western margin of the rønne graben in contrast to the marked inversion relative to the arnager block in the south-eastern part of the graben. b: note the very significant inversion of the kolobrzeg graben relative to the hanö bay basin. stina-1 is located in the kolobrzeg graben close to the koszalin fault. c: inversion of the kolobrzeg graben is also evident on the sw–ne section, east of bornholm. slightly modified from vejbæk et al. (1994). 626 a value of 0.48 %ro (ahlberg 1994). these data suggest burial depths between 2250 m and 3050 m. the bajocian–bathonian coal seams in the øresund-5 and øresund-7 wells situated immediately west of the inversion zone, were buried and later uplifted c. 1450 m and c. 1875 m, respectively, corresponding to modelled values of c. 1400 m for the øresund-5 well and c. 1500 m for the øresund-7 well (figs 10, 13; table 2). anholt well, kattegat the lower–middle jurassic strata encountered in the anholt well were buried and uplifted c. 975 m (fig. 10; table 2). the island of anholt is located on the boundary between the skagerrak–kattegat platform and the inversion zone. the uplift is mainly interpreted as being of neogene–pleistocene age, and the degree of uplift is slightly less than that inferred for the sæby-1 well (japsen 1993; michelsen & nielsen 1993). approximately 1000 m of uplift was estimated by nielsen et al. (2003, this volume) based on the same reflectance data but utilising a general north sea coalification curve. hydrothermal influence on the maturation of organic matter the vast majority of the reflectance values provide reasonable estimates of burial depths that are consistent i ii (a) top kimmeridgian 0.43 %ro 0.39 %ro 0.36 %ro 0.46 %ro 0.51 %ro katslösa mb döshult mb øresund-1 øresund-7 øresund-2 øresund-3 øresund-12 øresund-11 øresund-10 øresund-9 øresund-8 øresund-18 øresund-13 øresund-14 øresund-15 øresund-4 øresund-5 well section (b) 0 0.60 1.20 %r m et re s d ep th ( m ) 0 500 1000 1500 0 1600 3200 : measured %ro : modelled %r fig. 13. a: reconstructed stratigraphic columns at the location of the øresund-1 well. the reconstructions are based on the geological profiles along the tunnel transect (column i) and the bridge transect (column ii) shown on plate 21 in larsen et al. (1968). b: modelled reflectance values from the constructed pseudo-well compared to measured values. with stratigraphic thicknesses and other evidence of burial depths. however, the rank of the coals from the munkerup member, the uppermost bagå formation, and possibly the fuglunda member appear anomalously high. a mechanism other than thermal influence with burial depth is needed to explain these anomalies. the precursor peats of the uppermost bagå formation coals were formed in a relatively well-aerated environment, which may have produced hydrogen-poor and oxygen-rich huminite precursor material that matured at an enhanced rate (hao & chen 1992). however, this explanation cannot be applied to the coals of the munkerup member, and may not be sufficient to explain the large difference in reflectance between the coals from the sorthat and bagå formations on bornholm. therefore, it is proposed that the elevated maturity was the result of a local rise in temperature due to hydrothermal activity. hydrothermal processes associated with fault movements have been proposed to explain the occurrence of sideritic iron ores and copper mineralisation along the fyledal fault zone and zeolites in campanian deposits in the vomb trough (norling & bergström 1987; erlström 1994). middle jurassic volcanic activity occurred in skåne along fault and fracture zones, and hot brines influenced the diagenesis of rhaetian and lower jurassic sandstones and caused anomalously high vitrinite reflectances (klingspor 1976; norling & bergström 1987; ahlberg 1994; ahlberg & goldstein 1994). the silica-cemented höör sandstone (a lateral equivalent of the höganäs formation, fig. 2), which has not been buried more than a few hundred metres, was indurated at temperatures from 100–200°c, and the organic matter shows reflectance values that locally reach 0.9 %ro corresponding to a burial depth of c. 5700 m. bottomhole temperatures in the hans-1 well in the kattegat were high in lower jurassic strata at very shallow depths (45°c at 316 m) whereas normal temperatures were recorded at deeper levels. the real temperature of the jurassic deposits is even higher than indicated by the borehole measurement, because of the cooling effect of the drilling mud. the occurrence of hydrothermal activity along faults and the effect on the diagenesis of organic as well as siliciclastic deposits is thus well-documented. some of the observed anomalies in the jurassic coals may therefore be related to the middle jurassic volcanic activity in skåne. the uppermost bagå formation coals are situated close to the fault that separates the rønne graben from the crystalline basement of the bornholm high, and it is possible that hot formation waters expelled from deeply buried deposits percolated upwards along the fault and into the coal-bearing strata. the occurrence of abundant epigenetic pyrite in the coals and large pyrite nodules in sand interbeds lend support to this interpretation, as the most likely source of the sulphur is saline formation water from older sediments containing salt and gypsum (petersen et al. 2003, this volume). a similar situation has been reported from the carboniferous st. rose and chimney corner coalfields in nova scotia, canada (beaton et al. 1993). the munkerup member coals occur close to the northern and eastern bounding faults of the arnager– sose fault block, and these faults may also have functioned as conduits for hot formation waters, thus explaining the high rank of the coals. discussion and conclusions a well-constrained, uplift-corrected coalification curve has been constructed from reflectance data representing a wide range of burial depths of the upper triassic – lower cretaceous succession in the danish basin and the fennoscandian border zone. we suggest that this curve may be used as a standard with which new data can be compared. the reflectance values from wells in the fjerritslev trough and the terne-1 well deviate from the main coalification trend. basin modelling and velocity data from the chalk group (japsen 1998) suggest that the uplift correction may be erroneous due to overestimation of the burial depth based on fjerritslev formation sonic velocities in these wells. these data have thus not been incorporated in the standard curve. average reflectance values from the lower–middle jurassic coals in the fennoscandian border zone range from 0.29 %ro to 0.51 %ro (table 2), which are compatible with low rank coals, and directly indicate relatively shallow burial depths. the use of the regional curve to assess the burial depths and later uplift of coalbearing lower–middle jurassic strata in the fennoscandian border zone has provided reliable and consistent results (table 2). hydrothermal activity related to the middle jurassic volcanic event may locally have influenced the maturation of the coals in the hettangian–sinemurian of the höganäs basin, and possibly the coals in the fuglunda member at fyledal. the rank of the coals in the munkerup member (rønne formation) and uppermost bagå formation on bornholm was probably raised by hot formation waters expelled from deeply buried layers. 627 628 the reflectance data indicate c. 900–1400 m of postearly cretaceous uplift of the jurassic succession in the rønne–hasle fault block and the kolobrzeg graben (table 2). this degree of uplift is compatible with the amount of late cretaceous – early cenozoic inversion based on seismic interpretation (hamann 1994; vejbæk et al. 1994), and is similar to the 550–1000 m of inversion that has been interpreted from the sorgenfrei– tornquist zone further to the north-west (japsen 1993; michelsen & nielsen 1993). thus the total post-early cretaceous uplift in the bornholm area determined from reflectance values can be explained by inversion alone, without the need for significant neogene–pleistocene uplift. the data from the arnager–sose fault block indicate early jurassic burial of 625 m followed by a similar magnitude of middle–late jurassic uplift and erosion. post-early cretaceous uplift only amounts to 290 m. the estimated uplift of the jurassic section in the höganäs basin, reflecting both inversion and neogene– pleistocene uplift, is a minimum of 2010 m (table 2). this is similar to the total uplift of 1730–2000 m of the hans-1 well-section in the kattegat (japsen 1993; michelsen & nielsen 1993). the c. 1450–1875 m of uplift indicated by the øresund-5 and -7 data is a function of both neogene uplift and drag along the inversion zone. the rank of the bajocian fuglunda member coal bed suggests c. 1450 m of total uplift, which appears to be a reasonable estimate based on the likely overburden and regional evidence, but it cannot be precluded that the rank has been increased due to hydrothermal influence. the total uplift of the anholt well-section is estimated to be c. 975 m, rather than the > 1200 m suggested by jensen & michelsen (1992) but in agreement with the c. 1000 m of nielsen et al. (2003, this volume). the data from anholt and bornholm may suggest that the amount of neogene–pleistocene uplift in skåne was relatively small compared to the northern part of the danish basin and the fennoscandian border zone, but further data are needed for confirmation. acknowledgements ole v. vejbæk and peter japsen are thanked for discussions, jørgen a. bojesen-koefoed kindly assisted with data plotting, and niels-erik hamann generously allowed us to include an unpublished figure (fig. 7). the referees, lars n. jensen and john e.a. marshall, are thanked for their constructive comments. references ahlberg, a. 1994: diagenesis of rhaetian–hettangian coal-bearing siliciclastic strata in nw skåne, southern sweden. in: ahlberg, a.: deposition and diagenesis of the rhaetian– hettangian succession (triassic–jurassic) in southern sweden, 53 pp. unpublished ph.d. thesis, lund university, sweden. ahlberg, a. & goldstein, r.h. 1994: fluid inclusions in quartz overgrowths: evidence of mid-jurassic volcanic hot brine flushing at shallow burial depth in the lower jurassic höör sandstone, southern sweden. in: ahlberg, a.: deposition and diagenesis of the rhaetian–hettangian succession (triassic–jurassic) in southern sweden, 53 pp. unpublished ph.d. thesis, lund university, sweden. ahlberg, a., sivhed, u. & erlström, m. 2003: the jurassic of skåne, southern sweden. in: ineson, j.r. & surlyk, f. 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(eds): structural and tectonic modelling and its application to petroleum geology. norwegian petroleum society (npf) special publication 1, 153–162. peters, k.e. 1986: guidelines for evaluating petroleum source rock using programmed pyrolysis. american association of petroleum geologists bulletin 70, 318–329. petersen, h.i. 1994: depositional environments of coals and associated siliciclastic sediments in the lower and middle jurassic of denmark. the øresund-5, -7, -13, -15 and -18 wells. danmarks geologiske undersøgelse serie a 33, 55 pp. petersen, h.i. & nielsen, l.h. 1995: controls on peat accumulation and depositional environments of a coal-bearing coastal plain succession of a pull-apart basin; a petrographic, geochemical and sedimentological study, lower jurassic, denmark. international journal of coal geology 27, 99–129. petersen, h.i. & rosenberg, p. 1998: reflectance retardation (suppression) and source rock properties related to hydrogenenriched vitrinite in middle jurassic coals, danish north sea. journal of petroleum geology 21, 247–263. petersen, h.i., bojesen-koefoed, j.a. & nytoft, h.p. 1996: depositional environment and burial history of a lower cretaceous carbonaceous claystone, bornholm, denmark. bulletin of the geological society of denmark 43, 133–142. petersen, h.i., nielsen, l.h., koppelhus, e.b. & sørensen, h.s. 2003: early and middle jurassic mires of bornholm and the fennoscandian border zone: a comparison of depositional environments and vegetation. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 631–656 (this volume). radke, m., schaefer, r.g. & leythaeuser, d. 1980: composition of soluble organic matter in coals: relation to rank and liptinite fluorescence. geochimica et cosmochimica acta 44, 1787–1800. schmidt, b.j. 1985: a coal petrographic source rock evaluation of the rhaetic – jurassic – lower cretaceous sediments of the børglum-1 and uglev-1 wells, denmark. bulletin of the geological society of denmark 33, 239–252. schmidt, b.j. 1988: a source rock evaluation of the mesozoic sediments of the well hyllebjerg-1, danish subbasin. danmarks geologiske undersøgelse serie c 9, 105 pp. sorgenfrei, t. & buch, a. 1964: deep tests in denmark, 1935–1959. danmarks geologiske undersøgelse iii. række 36, 146 pp. suggate, r.p. 1998: relations between depth of burial, vitrinite reflectance and geothermal gradient. journal of petroleum geology 21, 5–32. surlyk, f., arndorff, l., hamann, n.-e., hamberg, l., johannessen, p.n., koppelhus, e.b., nielsen, l.h., noe-nygaard, n., pedersen, g.k. & petersen, h.i. 1995: high-resolution sequence stratigraphy of a hettangian–sinemurian paralic succession, bornholm, denmark. sedimentology 42, 323–354. sørensen, k. 1986: danish basin subsidence by cold triassic rifting on a lithospheric cooling background. nature 319, 660–663. taylor, g.h., teichmüller, m., davis, a., diessel, c.f.k., littke, r. & robert, p. 1998: organic petrology, 704 pp. berlin–stuttgart: gebrüder borntraeger. teichmüller, m. 1979: die diagenese der kohligen substanzen in den gesteinen des tertiärs und mesozoikums des mittleren oberrhein-grabens. fortschritte in der geologie von rheinland und westphalen 27, 19–49. thomsen, e. 1980: rank of coal and dispersed organic matter in rhaetian – jurassic – l. cretaceous deposits from the onshore part of the danish subbasin: interpretation and implications for the maturity of potential hydrocarbon source rocks, 269 pp. unpublished ph.d. thesis, university of aarhus, århus, denmark. thomsen, e. 1983: a coal petrographical investigation of the well farsø-1. dgu internal report (unnumbered), 14 pp. copenhagen: geological survey of denmark. thomsen, e. 1984: a coal petrographical investigation of the well års-1. dgu confidential report 1, 19 pp. copenhagen: geological survey of denmark. thomsen, e., lindgreen, h. & wrang, p. 1983: investigation on the source rock potential of denmark. geologie en mijnbouw 62, 221–239. thomsen, e., damtoft, k. & andersen, c. 1987: hydrocarbon plays in denmark outside the central trough. in: brooks, j. & glennie, k.w. (eds): petroleum geology of north west europe, 375–388. london: graham & trotman. vejbæk, o.v. 1985: seismic stratigraphy and tectonics of sedimentary basins around bornholm, southern baltic. danmarks geologiske undersøgelse serie a 8, 30 pp. vejbæk, o.v. 1989: effects of asthenospheric heat flow in basin modelling exemplified with the danish basin. earth and planetary science letters 95, 97–114. vejbæk, o.v. 1990: the horn graben, and its relationship to the oslo graben and the danish basin. in: neumann, e.-r. (ed.): rift zones in the continental crust of europe – geophysical, geological and geochemical evidence: oslo – horn graben. tectonophysics 178, 29–49. vejbæk, o.v., stouge, s. & poulsen, k.d. 1994: palaeozoic tectonic and sedimentary evolution and hydrocarbon prospectivity in the bornholm area. danmarks geologiske undersøgelse serie a 34, 23 pp. ziegler, p.a. 1982: geological atlas of western and central europe, 130 pp. the hague: elsevier for shell internationale petroleum maatschappij. 630 manuscript received 12 january 1996; revision accepted 5 march 1997. geological survey of denmark and greenland bulletin 38, 2017, 21-24 21 west of københavn, the top of the pre-quaternary limestone is found near the terrain surface. there is only a relatively thin cover of quaternary deposits, which makes the limestone vulnerable to pollution. region hovedstaden, being responsible for treating polluted sites, therefore asked geo and the geological survey of denmark and greenland to describe the geology and hydraulic characteristics of the limestone formations (galsgaard et al. 2014). during this work new information and data were collected and a revised geological model established for the area between københavn and roskilde (fig. 1). the model is based on seismic sections, a revised map of the pre-quaternary surface, biostratigraphy, borehole information and geophysical data. this paper presents the revised geological model. limestone formations the uppermost 100 m of the pre-quaterary deposits between københavn and roskilde comprises maastrichtian chalk, bryozoan limestone (stevns klint formation), the københavn kalk formation and the lellinge grønsand formation (figs 1, 2). cretaceous (maastrichtian) chalk is a carbonate mudstone. it is overlain by danian deposits in the area between roskilde and københavn, but occurs at the pre-quaternary surface in the southern part of the area. the early and middle danian bryozoan limestone, defined as the stevns klint formation by surlyk et al. (2006), usually contains 20 to 45% bryozoan fragments, but the formation also comprises mudstone or calcarenite with scattered bryozoans. usually the bryozoan limestone is deposited in mounds, which are strongly asymmetrical in the lower mound complex and less so in the middle and upper mound complexes (fig. 2). flint occurs as layers between the limestone layers and lenses of coral limestone also occur. the bryozoan limestone is 53–63 m thick in the københavn area (stenestad 1976). the københavn kalk formation is of late danian age. it is a sandy and silty carbonate mudstone defined by stenestad (1976). it has sub-horizontal layering with pronounced flint layers parallel to the layering. a log-stratigraphy was established in the københavn area by klitten et al. (1995; fig. 2). the formation is 40–45 m thick. the youngest formation, the lellinge grønsand formation, is of selandian age and consists of glauconite and carbonate-rich sand with layers and lenses of sandy limestone. glauconite-rich marl also occurs. structures and stratigraphy of danian limestone, eastern sjælland, denmark peter roll jakobsen, magnus marius rohde and emma sheldon a b c 1 2 3 4 7 65 biostratigraphical sample maastrichtian chalk stevns klint formation københavn kalk formation calcarenite lellinge grønsand formation danian selandian fault carlsberg fault roskilde fault søndersø valley ø resund fault risø fault ishøj fault zone 5 km roskilde københavn 5 kmk r 40 20 0 −20 −40 −60 −80 ka tri ne bje rg fa ult fig. 1. a: a revised geological map of the pre-quaternary surface in the københavn–roskilde area. the numbered dots show the location of analysed samples. r: roskilde. k: københavn. b: an east–west profile. c: location of the study area. © 2017 geus. geological survey of denmark and greenland bulletin 38, 21–24. open access: www.geus.dk/publications/bull 2222 biostratigraphy calcareous nannofossils are the main constituent of chalk deposits. calcareous nannofossils are thought to be the remains of the principal calcareous nannoplankton group: the haptophyte algae (bown & young 1998). they are useful as biostratigraphic markers as they have a widespread distribution and are present globally in the photic zone of almost all marine habitats. the nannofossil biostratigraphic dating of the outcrops and wells from eastern sjælland is based on the north sea zonation scheme of varol (1998) which, along with the danish onshore nannofossil zonation scheme of thomsen (1995) for comparison, is seen in fig. 2. the zonation scheme of varol (1998) was mainly based on north sea wells, but sections from onshore denmark were also used in its construction. as seen in fig. 2, the zonation scheme of varol (1998) allows for a higher resolution biostratigraphic breakdown than that of thomsen (1995). table 1 (youngest stratigraphy at the top) shows the biostratigraphic dating of samples taken from two outcrops and five wells (fig. 1). the results show that samples from the more easterly localities are the oldest (middle danian, subzones nntp2f–g), and the samples become younger to the west (upper danian–selandian subzones nntp4c and nntp5b). faults and interpreted faults the most pronounced tectonic feature in the subsurface of københavn is the se–nw-trending carlsberg fault that separates the københavn kalk formation from the stevns klint formation (fig. 1). the fault is one of a number of relay faults related to the tornquist–sorgenfrei wrench fault zone. the carlsberg fault can be regarded as a negative flower structure with a main offset between 50 and 100 m of the hanging-wall block down to the ne (fallesen 1995; jakobsen et al. 2002). the carlsberg fault may still be active as terrain movements have been detected across the fault (jakobsen et al. 2013) and neo-tectonic faulting of the quaternary cover is seen (kammann et al. 2016). the carlsberg fault is recorded on the seismic profile hgs002 (fig. 3) along with the øresund fault and the ishøj fault zone. the øresund fault is almost parallel to the carlsberg fault, but they merge about 12 km north of the fig. 2. chronostratigraphy, biostratigraphy, lithostratigraphy and lithology in the greater københavn area (modified from lund et al. 2002). numbered dots are the numbered localities in fig. 1a. * locality numbers refer to localities marked on figs 1, 2 table 1. biostratigraphical results 2 kallerup pit nntp4c early late danian 3 naverland 26, glostrup nntp2f middle danian 4 kirkebjerg parkvej 14 nntp2g middle danian 5 dalager 7 nntp2g middle danian 6 hesselager 17 nntp2g middle danian 7 hvidovre hospital nntp2f middle danian locality* nannofossil zone stage 1 hove limestone pit nntp5b-6 uppermost danian – selandian stage 1995 lithology se lan di an 9 nntp6 n n tp 5 f a b e lellinge grønsand formation 1 2 4 3 7 5 6 n n tp 1 b a n n tp 2 a b c d e n n tp 4 nntp3 a b c d f g d an ia n m aa st ric htia n ea rly m id dl e la te br yo zo an li m es to ne st ev ns k lin t f or m at io n kø be nh av n ka lk f or m at io n u pp er m ou nd c om pl ex m id dl e m ou nd c om pl ex u pp er m id dl e lo w er 1 2 3 4 5 6 7 8 ch al k lo w er m ou nd c om pl ex ca lc ar en iti c lim es to ne coral flint burrow hardground mounds and layers in bryozoan limestone thomsen varol 1998 23 seismic profile. the ishøj fault zone is a c. 5 km wide positive flower structure with folding between the faults and an over-all inversion across the zone. the inversion has caused the maastrichtian chalk to be present at the prequaternary surface in the ishøj area. one of the prominent continuous reflectors in the upper part of the seismic profile is outlined in fig. 3, and it dips c. 50 m from the ishøj fault zone to the eastern part of the area shown in the seismic section. the same dip is seen on the seismic section hgs-001 about 10 km north of hgs-002. the roskilde and risø faults are the westernmost faults presented in fig. 1; they form part of a series of n–s-trending relay faults in the roskilde area (pedersen & gravesen 2016). pre-quaternary surface the danian limestone forms the pre-quaternary surface in eastern sjælland and is the primary source of groundwater in the area. the large number of wells and well-logs and geological information recorded and stored in the national jupiter database allow a detailed mapping of the limestone surface. based on these well-logs and several geological models, a new map of the level of the pre-quaternary surface was constructed (fig. 4). in the northern part of the map area, the søndersø valley has a valley bottom c. 40–30 m below sea level and strikes sw–ne. several buried valleys have been identified and marked with red lines on the map (fig. 4). the longest of these is the herlev valley that is oriented parallel with the søndersø valley c. 5 km south to the of it. towards the west, the herlev valley almost intersects the smaller katrinebjerg valley. the general level of the prequaternary surface is 5–10 m higher south of the herlev and katrinebjerg valleys compared with the area between these valleys and the søndersø valley. two narrow valleys, the rådhus and vibenhus valleys, situated below the city of københavn are generally well mapped based on borehole data. several conspicuous depressions in the limestone surface are only identified from one or a few boreholes, e.g. in the northern part of taastrup and the eastern part of the herlev valley. it is possible that these small depressions are un-mapped valleys similar to the clear buried valleys identified on the map in fig. 4. geological model the revised geological map and model are presented in fig. 1. the map is based on interpretation of the map of the pre-quaternary surface, seismic sections and biostratigraphic analyses. the biostratigraphic analysis of the middle part of the stevns klint formation (fig 1) shows that the limestone in this area, at the pre-quaternary surhgs005 hgs003 500 0 tw ow ay tr av el ti m e (m s) hgs-002 hgs-002 4 km c arlsberg f ø resund f ishøj fz tw ow ay tr av el ti m e (s ) 0 0.5 w ehgs-005 hgs-003 hgs-001 hgs -005 hgs -003 fig. 3. seismic section hgs-002. the inset map shows the location of the profile and some additional profiles used for the interpretation. 30 to 40 20 to 30 10 to 20 0 to 10 −10 to 0 −20 to −10 −30 to −20 −40 to −30 −50 to −40 −60 to −50 > −60 burried valley elevation (m) 5 m contour 2.5 m contour 5 km n søndersø valley herlev valley vibenhus valleyrådhus valley katrinebjerg valley taastrup ishøj søndersø valley herlev valley vibenhus valleyrådhus valley katrinebjerg valley taastrup ishøjfig. 4. topography of the pre-quaternary surface in the københavn–roskilde area based on the latest data. 2424 face, is of middle danian age and belongs to the upper mound complex (fig. 2). compared with the depth to the maastrichtian–danian boundary, registered in adjacent boreholes, the thickness of the stevns klint formation is fairly uniform and comparable with thicknesses known from the københavn area, and a uniform thickness of the limestone unit is consequently assumed. previous maps do not show the presence of the københavn kalk formation between the lellinge grønsand formation and the stevns klint formation in the western part of the area (stenestad 1976). with a dip of c. 50 m from the ishøj fault zone to roskilde, it is therefore expected that the københavn kalk formation is present between the stevns klint formation and the lellinge grønsand formation. this was confirmed by biostratigraphical analyses of samples from localities 1 and 2 (figs 1, 2). however, the two samples representing københavn kalk formation are from the uppermost and lowermost parts of the formation, respectively. this means that the katrinebjerg fault (fig.1), interpreted from the pre-quaternary surface map, is in fact a fault, with an offset of c. 50 m. at hove (locality 1, fig. 1) a geophysical log from borehole dgu 200.5248 shows the same log pattern as it is known from the københavn area and shows the presence of the whole formation. many samples from boreholes within the western københavn kalk formation area are described as bryozoan limestone. this could represent a facies variation with a higher concentration of bryozoans than is usually seen in the københavn area. however, it is also possible that the tectonic conditions are more complex, with more faults than those documented so far. conclusions the geological map of the pre-quaternary surface is revised for the area between københavn and roskilde. the revision is based on new seismic sections, a map of the prequaternary surface and biostratigraphic analyses. the ishøj fault zone is a positive flower structure, and a prominent structural feature in this area. between the ishøj fault zone and the roskilde fault, the limestone dips westwards. the upper danian københavn kalk formation is documented in the western part of the pre-quaternary surface between københavn and roskilde and may have a higher content of bryozoans than in the københavn area. acknowledgements the project on geological and hydrological knowledge acquisition on danian limestones between københavn and roskilde was financed by region hovedstaden. region hovedstaden kindly allowed us to use the results from biostratgraphical analyses conducted during various projects. references bown, p. & young, j. 1998: introduction. in: bown, p.r. (ed.): calcareous nannofossil biostratigraphy. british micropalaeontological society publication series, 1–15. fallesen, j. 1995: stratigraphy and structure of the danian limestone on amager, examined with geophysical investigations – especial with regard to the carlsberg fault. unpublished msc thesis, university of copenhagen. galsgaard, j., rohde, m., jakobsen, p.r. & jakobsen, r. 2014: strømning og stoftransport i kalklagene på den københavnske vestegn. geologisk og hydrologisk vidensopsamling og typemodel, 97 pp. unpublished report, geo projekt nr. 37208, rapport 1, 2014-06-19. jakobsen, p.r., fallesen, j. & knudsen, c. 2002: strukturer i den københavnske undergrund – folder, forkastninger og sprækker. in: frederiksen, j.k et al. (eds): ingeniørgeologiske forhold i københavn. danish geotechnical society bulletin 19, 19–29. jakobsen, p.r., wegmuller u., capes r. & pedersen s.a.s. 2013: terrain subsidence detected by satellite radar scanning of the copenhagen area, denmark, and its relation to the tectonic framework. geological survey of denmark and greenland bulletin 28, 25–28. kammann, j., huebscher, c., boldreel, l.o. & nielsen, l. 2016: highresolution shear-wave seismics across the carlsberg fault zone south of copenhagen: implications for linking mesozoic and late pleistocene structures. tectonophysics 682, 56–64. klitten, k., ploug, c. & olsen, h. 1995: geophysical log-stratigraphy of the københavn limestone. danish geotechnical society bulletin 11(5), 5127–5134. lund, n.s., nielsen, l.h. & knudsen, c. 2002: københavns undergrund med fokus på danien aflejringerne. in: frederiksen, j.k. et al. (eds): ingeniørgeologiske forhold i københavn. danish geotechnical society bulletin 19, 5–18 pedersen, s.a.s. & gravesen, p. 2016: tectonic control on the formation of roskilde fjord, central sjælland, denmark. geological survey of denmark and greenland bulletin 35, 35–38. stenestad, e. 1976: københavnsområdets geologi især baseret på citybaneundersøgelserne. danmarks geologiske undersøgelse iii. række 45, 149 pp. surlyk, f., damholt, t. & bjerager, m. 2006: stevns klint, denmark: uppermost maastrichtian chalk, cretaceous–tertiary boundary and lower danian bryozoan mound complex. bulletin of the geological society of denmark 54, 1–48. thomsen, e. 1995: kalk og kridt i den danske undergrund. in: nielsen, o.b. (ed.): danmarks geologi fra kridt til i dag. aarhus geokompendier 1, 31–67. varol, o. 1998: palaeogene. in: bown, p.r. (ed.): calcareous nannofossil biostratigraphy. british micropalaeontological society publication series, 200–224. authors’ addresses p.r.j. & e.s.,geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: prj@geus.dk m.m.r., geo, maglebjergvej 1, dk-2800 kgs. lyngby, denmark. e2019430303-01 characterisation of incinerator bottom ash from a danish waste-to-energy plant: a step towards closing the material cycle rune j. clausen*1, per kalvig1 and jonas nedenskov2 geus bulletin is free to individuals and institutions in electronic form. the author(s) retain copyright over the article contents. research article | open access geus bulletin vol 43 | e2019430303 | published online: 20 december 2019 https://doi.org/10.34194/geusb-201943-03-03 cineration process and these are typically composed of silicarich melts with recognisable remnants of glassware, ceramics and metal fragments. the extreme heterogeneity among iba fragments is reflected not only in their chemical composition, but also in their variable density, shape and size from 2000 mm to less than 1 mm (fig. 1). iba characteristics, including their physical form and chemical composition, are the result of a wide range of dynamic parameters. first, the composition of the input to the incinerators – the municipal solid waste – is physically and chemically heterogeneous and complex. further, the temperature and fugacity conditions in the oven chamber vary locally due to the design of the chamber as well as irregular local flux-effects, generated by variations in the solid-waste input. all in all, iba should be considered a multi-commodity deposit, for which detailed characterisation of the constituent materials is key to revealing its resource potential. methods four iba samples were collected in duplicate from amager bakke plant each day for 30 consecutive days in november 2017 (series a–d, fig. 2). series d was preserved as an archive sample, and a, b and c were analysed for their composition. each sample weighed c. 15 kg, and the total sampled material weighed c. 1800 kg. representative sampling and physical characterisation representative sampling of such heterogeneous material is a serious challenge. for example, the iba is moving on a shaking belt and is not distributed evenly, iba is hot (around 50–100°c) and fragments can be very large. we attempted to overcome the challenges by sampling a cross-section with a steel shovel taking care to avoid subconsciously ‘fishing’ for the un sustainable development goal 12, regarding responsible production and consumption of raw materials, guides ongoing international efforts to enhance sustainability in all parts of the mineral sector. of particular interest, is improving the recyclability of secondary waste streams and thereby increasing the efficiency of recycling end-oflife products. municipal solid waste – residual waste from household and industry – constitutes one of these secondary streams. it is typically incinerated in waste-to-energy plants producing two types of waste streams that carry a raw material resource potential: incinerator bottom ash (iba) and incinerator fly ash (ifa). iba is of particular interest in the recycling industry, where it is commonly recycled to produce three main fractions: (i) ferrous material, (ii) non-ferrous material, and (iii) residual slag. in most cases the two metal fractions are separated further downstream in the value chain, prior to smelting. the residual, non-magnetic fraction (typically 0–45 mm) is used mainly as construction aggregate. improvements in the efficiency of existing separation technologies are still being made, but less effort is focussed on characterising the fundamental composition and mineral resource potential of iba. for this reason, the urban-x project was launched by the geological survey of denmark and greenland (geus) to characterise the composition and resource potential of various waste streams at amager bakke waste-to-energy plant in copenhagen, denmark. this paper discusses some of the main outcomes of the urban-x project with respect to iba, and a full analysis of all waste streams analysed at amager bakke is available in clausen et al. 2019. incinerator bottom ash iba material consists of all the non-combustible components of domestic and industrial waste. it can include components such as construction steel, household glass fragments, toys, electronic devices, cans, batteries, sofa springs, and boulder size aggregates. some of the aggregates form during the inhttps://doi.org/10.34194/geusb-201943-03-03 e2019430303-02 particular pieces. fragments of up to c. 200 mm were sampled, though due to the sampling bias associated with large fragments (relative to the size of the sampler) among other things, only fragment sizes less than or equal to 63 mm were considered representatively sampled. to create four sample series (a–d) carrying the same 30-day representation, four samples was extracted immediately one after the other. for physical characterisation, we combined one series into a composite sample, which represented the average iba material collected during the 30 days of sampling (c series, fig 2). first, all fragments larger than 63 mm were removed by hand and ruler in preparation for the splitter, where each sample was divided to retrieve a representative subsample. combining all 30 subsamples of series c produced a composite sample of 17 kg. the composite sample was then sieved into seven size fractions, which were further separated to isolate a number of characteristic components: ferro magnetic metal, nonmagnetic metal, glass, ceramics and building aggregates and melt. glass, ceramics and building aggregates were identified by manual-visual sorting. ferro magnetic metal was captured by a magnet and non-magnetic metal was identified with a metal-detector. the remaining material was classed as predominantly melt – fragments or conglomerations of fragments, which were partly or entirely melted, and did not belong to the other categories. the resulting fractions and material classes were weighed. all metal fragments underwent further analysis with x-ray fluorescence (xrf) to categorise them according to alloying elements; non-magnetics were subdivided into aluminium, alloyed aluminium, copper alloyed copper, and other metals. additional subdivisions were made according to the degree of degradation; glass for example was subdivided into four sub-classes. this part of the study is not reported in this article but can be found in clausen et al. 2019. chemical characterisation a classical mineral exploration approach would be to produce an iba ‘whole-rock’ chemical signature to identify potentially economic elements and minerals. but this is neither possible nor meaningful for iba material for the following reasons: 1. the technical challenge: homogenisation by means of crushing and milling of iba material is expensive, if not impossible, due to the content of ductile metal fragments. 2. the distributional challenge: the ‘whole rock’ chemical signature represents elements hosted by myriad chemically different fragments of materials and melts. however, it does not reveal the extent to which the element is available to mining/recycling, since it does not describe the size, shape and elemental composition of the particular fragments it is associated with. it follows that neither the ‘ore grade’ nor the ‘ore value’ of iba can be established only on the basis of bulk geochemical data. however, chemical data may point to elements occurring in elevated concentrations, on which further studies are required in order to assess their economic potential. this approach was applied in the urban-x project where a total of 62 1 cm b 1 cm c a fig. 1: example waste from amager bakke waste-to-energy plant. a: unprocessed incinerator bottom ash (reproduced from clausen et al. 2019). b: non-deformed glass fragments (sample number 567662 a; clausen et al. 2019). c: magnetic, non-deformed metal (sample number 567662 a; clausen et al. 2019). chemical elements were analysed by icp-ms and icp-oes on each of the 30 iba samples (the 0–63 mm size fraction from the b series after removing magnetic metal; fig 2). samples were prepared and analysed by actlabs, canada. a full description of sample preparation and analytical methods is supplied as supplementary information (file s1; https://doi. org/10.34194/geusb-201943-03-03). results and discussion five material classes dominated the coarse fraction of the iba-bulk sample: magnetic metal (29 wt%), non-magnetic metal (6 wt%), glass (14 wt%), ceramics and building aggregates (14 wt%) and melt (37 wt%; figs 1b, c). the composite iba-bulk sample consisted of 66 wt% coarse (2–63 mm) material and 34 wt% fine (< 2 mm) material. comparing the physical characterisation of each sieved fraction, it follows that each material class correlates to some extent with fragment size (fig 3). for example, building aggregates and magnetic metals are concentrated in the >32 mm fractions, while glass and non-magnetic metal are concentrated in the <32 mm fractions. distribution of the non-magnetic fraction is however less distinct across the various material classes (fig. 3). the annual resource potential of the five main classes in the 2–200 mm material fraction from the amager bakke plant are estimated and presented in table 1. the total mega tons per annum (mtpa) iba production in denmark and the eu-28 is 0.6 mtpa (miljøstyrelsen 2016) and 16 mtpa (iswa 2015), respectively. assuming the distribution of the five material groups measured for amager bakke iba is representative of iba in denmark and the eu, we can apply these class distributions to the national and eu-level data (table 1). for example, magneticand non-ferrous scrap could be as high as 230 000 tpa and 50 000 tpa, respectively, in denmark, and 4 200 000 tpa and 1 000 000 tpa, respectively, in the eu-28 (table 1). this eu-28 estimate must, physical characterisation >63 mm 31.5–63 mm 16–31.5 mm 8–16 mm 4–8 mm 2–4 mm >2 mm <2 mm sievingsplitterc series icp-ms archive archive archive d series b series grain size distribution chemical characterisation composite samples sievinga series archive fig. 2. sampling workflow and methods applied to each series (a–d) of sampled iba (modified from clausen et al. 2019). e2019430303-03 80 000 1 800 000 1 700 000 910 000 3 600 000 4 700 000 2 400 000 16 000 000e 30 000 68 000 63 000 34 000 137 000 177 000 91 000 600 000d 4000 9100 8400 4600 18 000 24 000 12 000 80 000 5b 11 10 6 23 29 15c 100 14 14 6 29 37 100 ceramics and aggregates glass non-ferrous metal magnetic metal (steel) slag melt total 0–2 mm fraction 2–63 mm >63 mm (mostly magnetic metal) estimated volume (tpa) all wte incinerators eu-28 estimated volume (tpa) all wte incinerators denmark estimated volume (tpa) adjusted (wt%)a measured (wt%) amager bakke a measured content adjusted to include estimates of the >63 mm and 0–2 mm fractions. b rough estimate of the >63 mm fragments in iba from amager bakke. c production from amager bakke, as estimated by amager resource center. d iba production in denmark based on multiple sources (miljøstyrelsen 2016; dansk affaldsforening et al. 2016). e iba production in the eu-28 estimated by iswa (2015). table 1. estimated annual resource volume produced by amager bakke waste-to-energy (wte) plant. potential volume in denmark and the eu-28 is extrapolated from the distribution among raw material groups observed at amager bakke https://doi.org/10.34194/geusb-201943-03-03 https://doi.org/10.34194/geusb-201943-03-03 however, be considered speculative and further studies are needed to quantify the resource potential in the eu. complete chemical analysis of the 0–63 mm iba fraction can be found in clausen et al. (2019). here we present only the nine elements that occurred in concentrations 10 times higher than the average crustal concentration and are thus considered candidates for resource extraction (table 2). when ranking these elements according to their level of enrichment compared to crustal concentration, sb, au and pb rank highest. since the various classes of potentially economic materials are unevenly distributed throughout the iba size fractions, it follows that the chemical composition also varies with grain size. therefore, a relevant question is, where in the iba are the nine elements in table 2 concentrated, i.e. elevated above the average (0–63 mm fraction) concentration? the 2–63 mm fraction of melt fragments were not measured in this study, but similar iba material (20–40 mm size fraction of melt fragments) was analysed at amager forbrænding in 2015 (kalvig et al. 2016; in service between 1970 and c. 2017, after which the new amager bakke plant operates in its place). the 20–40 mm melt fraction at amager forbrænding has a similar composition to the 0–63 mm fraction at amager bakke (table 2). in general, the 2–63 mm non-melt fragments (i.e. glass, building materials, magnetic metals and non-magnetic metal) originates from products used in household and industry, in which the elements in table 2 would not serve a functional purpose and are undesirable for economic, health and environmental reasons. measurements with xrf were carried out to see if any of the elements in table 2 could be detected among each category of non-melt fragments. none of the elements, except for zinc, registered above the limit of detection. zinc only registered in measurable quantities in a few non-ferrous metal fragments (data not shown). the xrf measurements thus support our general assumption, that the nine elements are mostly absent in the 2–63 mm non-melt fragments. we can thus assume that the 0–2 mm fraction is a source of elevated concentrations of the nine elements of interest. previous studies in holland support the notion that heavy metals are concentrated in the iba fines (muchova et al. 2009; muchová & rem 2006). summary and outlook iba flows at the amager bakke plant likely carry a secondary raw material potential that is not yet fully realised. these include: (1) glass, (2) increased recycling efficiency on ferrous and non-ferrous metals, (3) higher value usage of melts and ceramics e.g. in concrete and asphalt and (4) potential extrac0 20 40 60 80 100 2–4 4–8 8–16 16–32 32–63 63–200 size fraction (mm) c om po sit io n of 5 m at er ial cla ss es (% ) melt ceramics & building aggregate glass non-magnetic metals magnetic metals fig. 3. the distribution of main components according to fragment size. the characterized 2–63 mm fractions were considered representative of the total iba material. the 63–200 mm was not considered representative due to its size relative to the sampler, but a rough characterisation using a magnet was made nonetheless on the sampled material (modified from clausen et al. 2019). e2019430303-04 52 330 866 1.5 312 5.6 4.1 3355 118 210 154 98 72 66 41 31 39 12 42 480 980 1.8 150 3.2 4.7 3100 74 0.2 3.1 10 0.025 2.2 0.08 0.15 79 6.3 ppm ppb ppm ppm ppm ppm ppm ppm ppb sb au pb bi sn ag cd zn pd measured ave. conc.c earth crust enrichment factorb (no units) 0–63 mm (amager bakke) 20–40 mm (melt; amager forbrænding) earth crust ave. conc.a measured ave. conc. a obtained from periodictable.com. b earth crust enrichment factor is the element concentration relative to crustal concentration. table 2. content of selected elements in incinerator bottom ash produced by the amager bakke waste-to-energy plant c measured concentration of 10 fragments of melt materials from amager forbrænding, 2015 (kalvig et al. 2016). tion of metals such as antimony, gold, lead and zinc from the 0–2 mm fraction. future work should characterise these resource potentials in greater detail to enable their commercial exploitation. in addition, a comprehensive physical and chemical characterisation of the iba at similar plants would be a powerful tool to characterise the drivers of iba resource potential and to provide a foundation for closing some of the major gaps in the circular economy. to further enable commercial exploitation, efforts should be made to develop a preparatory recycling method of iba flows. a preparatory method could for example involve a washing phase, as part of the existing water-cooling basin at amager bakke – a method whereby clean fragments larger than 2 mm are removed in one material stream, and fines undergo a density separation to separate a heavy density concentrate. this method could (1) enable recycling of clean, non-metal fragments due to the absence of heavy metals – and make the fragments available for visual sorting techniques, (2) increase performance of existing mechanical sorting techniques applied to metal fragments, (3) enable recovery of antimony, gold, lead and zinc from the heavy concentrate, (4) save storage space currently used to store iba for carbonation (whereby heavy elements are stabilised) – which in turn would minimise the oxidation of stored metals. references clausen r., kalvig, p. & nedenskov, j. 2019: karakterisering af slagge og flyveaske fra affaldsforbrændingsanlægget amager bakke. overvejelser om råstofpotentialet. 2. udgave. mima rapport 2019/1, 231 pp. videncenter for mineralske råstoffer og materialer (mima) and geological survey of denmark and greenland (geus), denmark. accessed november 2019 at http://mima.geus.dk/wp-content/uploads/karakerisering-af-slagge-og-f lyveaske-fra-amager-bakke-overvejelser-omr%c3%a5stofpotentialet-2.-udgave-clausen-et-al-2019.pdf dansk affaldsforening et al. 2016: beate, benchmarking af affaldssektoren 2016 (data fra 2015), forbrænding, 29 pp. the danish energy agency, denmark. accessed november 2019 at https://ens.dk/sites/ens.dk/files/ affald/beate_afrapportering_forbraending_2016_29maj2017.pdf iswa 2015: bottom ash from wte plants – metal recovery and utilization. unpublished report, international solid waste association (iswa), austria. kalvig, p., clausen, r. & nedenskov, j. 2016: karakterisering af slagge og røggasaffald fra amager ressource center, 161 pp. unpublished report. videncenter for mineralske råstoffer og materialer, denmark. miljøstyrelsen 2018: affaldsstatistik 2016, 66 pp. the danish environmental protection agency, denmark. miljøprojekt no. 2020. https:// www2.mst.dk/udgiv/publikationer/2018/06/978-87-93710-39-9.pdf miljøstyrelsen 2016: kommissionens forslag til ændring af bilag iii i affaldsdirektivet (2008/98ef) med hensyn til fareegenskaber hp 14 (økotoksisk) (komitesag). notat til folketingets europaudvalg. ref. thfru. https://www.ft.dk/samling/20161/almdel/mof/bilag/35/1678142/index.htm muchová, l. & rem, p. c. 2006: metal content and recovery of mswi ash in amsterdam. in: v. popov, v. et al. waste management and the environment iii. wit transactions on ecology and the environment, 92. 211–216. https://doi.org/10.2495/wm060231 muchova, l., bakker, e. & rem, p. 2009: precious metals in municipal solid waste incineration bottom ash. water air soil pollution: focus 9, 107–116. https://doi.org/10.1007/s11267-008-9191-9 acknowledgements our thanks go to the reviewers tod waight and gang liu for their comments, and to marija blazanovic, tonny b. thomsen, nynke keulen, sebastian næsby malkki and frederik tevil for their help and guidance. the urban-x project was funded by amager bakke and videncenter for mineralske råstoffer og materialer (mima). how to cite clausen, r.j., kalvig, p. & nedenskov, j. 2019: characterisation of incinerator bottom ash from a danish waste-to-energy plant: a step towards closing the material cycle. geological survey of denmark and greenland bulletin 43, e2019430303. https://doi.org/10.34194/ geusb-201943-03-03 *corresponding author: rune clausen | e-mail: rjc@geus.dk 1 geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350, copenhagen k, denmark. 2 amager ressource center, vindmøllevej 6, dk-2300, copenhagen s, denmark. e2019430303-05 http://mima.geus.dk/wp-content/uploads/karakerisering-af-slagge-og-flyveaske-fra-amager-bakke-overvejelser-om-r%c3%a5stofpotentialet-2.-udgave-clausen-et-al-2019.pdf http://mima.geus.dk/wp-content/uploads/karakerisering-af-slagge-og-flyveaske-fra-amager-bakke-overvejelser-om-r%c3%a5stofpotentialet-2.-udgave-clausen-et-al-2019.pdf http://mima.geus.dk/wp-content/uploads/karakerisering-af-slagge-og-flyveaske-fra-amager-bakke-overvejelser-om-r%c3%a5stofpotentialet-2.-udgave-clausen-et-al-2019.pdf https://ens.dk/sites/ens.dk/files/affald/beate_afrapportering_forbraending_2016_29maj2017.pdf https://ens.dk/sites/ens.dk/files/affald/beate_afrapportering_forbraending_2016_29maj2017.pdf https://www2.mst.dk/udgiv/publikationer/2018/06/978-87-93710-39-9.pdf https://www2.mst.dk/udgiv/publikationer/2018/06/978-87-93710-39-9.pdf https://www.ft.dk/samling/20161/almdel/mof/bilag/35/1678142/index.htm https://www.ft.dk/samling/20161/almdel/mof/bilag/35/1678142/index.htm https://doi.org/10.2495/wm060231 https://doi.org/10.1007/s11267-008-9191-9 https://doi.org/10.34194/geusb-201943-03-03 https://doi.org/10.34194/geusb-201943-03-03 mailto:rjc%40geus.dk?subject= _goback geological survey of denmark and greenland bulletin 33, 2015, 25-28 25 acoustic events on a small seismological network – shock waves from thunder and fireballs peter h. voss, trine dahl-jensen and tine b. larsen th e geological survey of denmark and greenland (geus) operates a network of seismic stations in denmark primarily to detect earthquakes. but from time to time other sources than earthquakes generate seismic signals that are detected at the stations. here we show that both meteoroids and thunder have generated seismic signals with high signal-to-noise ratios at some of geus’ seismic stations (fig. 1). th e purpose of the seismic stations is to provide data for the earthquake database of the kingdom of denmark, hosted and maintained by geus. in order to avoid that the earthquake database is contaminated by other events not related to tectonism, these events are given special markers when possible. meteoroids in their fall through the atmosphere two meteoroids generated sonic signals close to danish seismic stations with suffi cient energy to cause seismic signals (fig. 2). fireballs of the meteoroids were observed. th e sonic signal is interpreted as the sonic shock wave that is generated when the speed of the fi reball exceeds the speed of sound. th e fi rst observation was made on 17 january 2009 when a meteoroid entered the atmosphere from space above the baltic sea and landed fig. 1. map of denmark showing the location of seismic stations and towns mentioned in the text. the arrows show the paths of fireballs. 10°e 55° 57°n 50 km denmark sweden læsø aalborg kattegat germany stevns maribo frederikshavn klokkerholm baltic sea seismic station town or village fig. 2. a: seismograms from the two seismic sensors that recorded the passing shock wave of the 2009 fireball. data have been band-pass filtered from 5 to 10 hz. b: seismograms from three seismic sensors that recorded the shock wave of the 2014 fireball showing the first motion downward on three stations. data have not been filtered. c: the arrows mark the pulses of the n-wave of the 2014 fireball. data were band-pass filtered from 1 to 10 hz. b c station 2,z station 4,z station 5,z 1 2 s 1 2 s lld 2,z station 2,z station 4,z 1 2 3 4 5 6 7 8 9 s a lld 1,z © 2015 geus. geological survey of denmark and greenland bulletin 33, 25–28. open access: www.geus.dk/publications/bull 2626 near the town of maribo, lolland (fig. 1), where a meteorite with a weight of 25.8 g was found (haack 2012; haack et al. 2012). th e second meteoroid was recorded on 16 june 2014 on a small network of six seismic stations north of aalborg, jylland. th is meteoroid entered the atmosphere south of frederikshavn and possibly ended in the kattegat south of the island of læsø (fig. 1). no meteorite was found. th e seismograms of the shock waves from the 2009 fi reball are seen in fig. 2a. th ey were recorded on the geus seismic station located on stevns. th e station is equipped with two vertical sensors installed 208 metres apart, with sensor lld2 located 150° south-east by south of sensor lld1. th e acoustic signal arrived 0.33 s later at lld2 than at lld1 which suggests that the source of the signal was located to the south-east of the stations. th e duration of the signals diff ers by a factor of four, which might be an eff ect of the diff erent conditions at the sites where the sensors are installed. th e lld1 signal is from the original sensor, located in a small vault dug into the fi eld. th e lld2 signal is from an experimental installation of a similar sensor next to the wall of a large barn. th e longer duration of this signal was probably caused by the resonance of the shock wave by the barn. when seismic data from the 2014 fi reball were fi rst analysed, the lack of s-wave energy suggested that the source of the signals was a mine explosion in kattegat. but closer inspection of the signals showed that the fi rst motion of the detected p-wave was downward, consistent with dilatation at the source (fig 2b). th e expected fi rst motion of the p-wave from an explosion is upward in response to the compression exerted by the source on the surroundings. an analysis of the signals recorded on the six stations, seen in fig. 3, shows that the shock wave arrived at the stations from an almost easterly direction of approximately 96°, with an apparent surface velocity of 1255.8 m/s. assuming a sound speed of 342.2 m/s the incidence angle to the seismic stations is 15.9° from vertical. th ese results fi t well with the fi reball observations published at http://stjerneskud.info. th e observed fi reball trajectory derived by sørensen (2014) based on phonorth south east west fig. 3. seismic recordings from six seismic stations of the shock wave caused by the fireball on 16 june 2014. data are the up/down z component of the data. amplitudes are normalised and the data were band-pass filtered from 5 to 10 hz. the time is utc. fig. 4. photograph showing the meteoroid trajectory. the photograph was taken with a specially designed fireball camera on 16 june 2014 at 1:15 p.m. local time, in klokkerholm, jylland. the camera was turned upward and equipped with a fisheye objective. the meteor was visible for 3.9 s. the image was created by stacking all video frames showing the fireball. courtesy of kim lang, klokkerholm and anton sørensen, http:// stjerneskud.info. station 1,z station 2,z station 3,z station 4,z station 6,z station 5,z 23:19:00 23:19:1023:19:05 27 tographs taken at two locations in northern jylland (fig. 4) corresponds well with our calculation of the azimuth angle. th e crossing of the azimuth and the trajectory suggest that the source of the shock wave came from an altitude of 46 km in the upper part of the stratosphere over the island of læsø. but since sound in the atmosphere oft en follows nonlinear ray paths as seen by the low incidence angle, the derived altitude is very uncertain. th e changes in air pressure in the shock wave generate a seismic p-wave in the upper part of the earth. since the seismic p-wave travels faster than the shock wave, it can sometimes be detected just before the arrival of the shock wave as shown by, e.g. kanamori et al. (1991). but in the data from the two fi reballs that we have observed, we see no indication of p-wave signals arriving before the shock wave. as the sonic boom travels through the atmosphere, the change in pressure has an n-shaped pulse. at the conversion to seismic energy, the n-wave shows as two pulses on the seismogram (kanamori et al. 1991; cates & sturtevant 2002). th ree examples of n-waves recorded aft er the 2014 fi reball are shown in fig. 2c. th ese should not be confused with p-waves, which arrive through the ground. thunder th e six seismic stations north of aalborg, installed in a network with a radius of 5 km, have made it possible to detect signals that are normally regarded as noise, for instance thunder. th under couples to the ground like the sonic booms from, e.g. meteoroids, but since lightning oft en covers very large areas and occurs in sequences, the seismic signal consists of many peaks and is less impulsive (e.g. kappus & vernon 1991). an example of at least seven thunder signals within a short time window is seen in fig. 5. th e distance to a thunderstorm that will generate an observable signal depends on the power of the acoustic signal released by the thunder and the composition of the atmosphere at the time of the thunder, since changes in the atmosphere can dampen or amplify acoustic signals in diff erent directions. most of the thunder we have observed occurred close to our seismic stations, but we have also observed thunder up to 30 km from a seismic station. th e thunder signals diff er from earthquake signals and explosions by the absence of body waves and are characterised by an apparent surface velocity around the speed of sound. it is the apparent surface velocity of the thunder signals that discriminates the seismic signals from other noise signals, such as traffi c. since the distance between seismic stations is usually long (50 km), thunder will most oft en only be recorded by a single seismic station, whereas a minimum of three stations is needed to estimate the apparent surface velocity. it is therefore normally not possible to positively identify sonic waves. in order to identify thunder signals we verify the observations with the lightning measurements performed by the danish meteorological institute. maps of observed lightning in denmark are presented on the institute’s webpage (http://www.dmi.dk/vejr/maalinger/lyn/). discussion th e signals caused by sonic booms only constitute a small fraction of the observed seismic events in the geus earthquake database. but when they occur they must be identifi ed so that they do not contaminate the seismic data. th e sonic boom from the fi reball on 16 june 2014 was not reported by any persons. th e seismic recordings only contain signals u p to 50 hz which is close to the lower limit of the human hearing range (on average 31 hz). th e frequency spectra in fig. 6 show that the recorded seismic signal of the meteoroid shock wave on average contained less energy above 31 hz than for instance a thunder signal recorded on 1 november 2014. at around 15 hz the two signals have similar amplitudes, but again at lower frequencies the thunder signal is on average higher. with an energy content lower than a thunder signal, the sonic boom was probably not audible to humans. figure 6 also shows the spectra of the p-wave of a local magnitude 1.2 earthquake recorded on 4 august 2014, approximately 39 km from the seismic station. th e spectra of the p-wave are station 3,z station 5,z 11:25 11:26 11:27 11:28 station 1,z fig. 5. seismograms showing 4 minutes of data from the up/down sensor recorded at three stations as a thunderstorm passed by on 1 july, 2014. at least seven signals caused by thunder are seen. time is utc, data were band-pass filtered from 10 to 30 hz. www.dmi.dk/vejr/maalinger/lyn/ 2828 comparable to the thunder signal, at frequencies lower than 40 hz. hearing such a signal would require a coupling of the energy from the ground to the air, which would introduce an energy loss. when people send geus reports on earthquakes sounds, they are oft en related to movements in buildings. geus did not receive any reports from this earthquake being felt or heard. it is oft en the ground impact of meteors that is associated with danger, but the meteor that hit chelyabinsk, russia, on 15 february 2013 was a clear reminder that the shock wave may pose a signifi cant risk. more than 1600 people were hurt from falling debris and over 7300 buildings were damaged. seismic waves were observed at distances of more than 4000 km. th is corresponds to an earthquake magnitude of 3.6 (heimann et al. 2013). th e challenge of marking seismic events of non-tectonic nature in the earthquake database remains and we are still not able to identify and mark them all. although we have recorded explosions, glacial earthquakes, thunder and fi reballs we still have not detected any footquakes in denmark. footquakes are seismic signals from sport events like football, an example comes from the 2006 african cup. when cameroon scored goals during the games, people in cameroon who watched the games on tv jumped a lot and generated simultaneous signals on 20 seismic stations across the country (euler 2007). references cates, j.e. & sturtevant, b. 2002: seismic detection of sonic booms. journal of the acoustical society of america 111, 614–628. euler, g.g., wiens, d.a. & loft on, k.m. 2007: footquakes. iris newsletter 1, p. 13. haack, h. 2012: meteoritter – tidskapsler fra solsystemets oprindelse, 189 pp. copenhagen: gyldendal. haack, h. et al. 2012: maribo – a new cm fall from denmark. meteoritics and planetary science 47, 30–50. heimann, s., gonzález, á., wang, r., cesca, s. & dahm, t. 2013: seismic characterization of the chelyabinsk meteor’s terminal explosion. seismological research letters 84, 1021–1025. kanamori, h., mori, j., anderson, d.l. & heaton, t.h. 1991: seismic excitation by the space shuttle columbia. nature 349, 781–782. kappus, m.e. & vernon, f.l. 1991: acoustic signature of thunder from seismic records. journal of geophysical research-atmospheres 96, 10989–11006. sørensen, a.n. 2014: http://www.stjerneskud.info/fi reball/ event2014-06-16-01-15/ authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: pv@geus.dk 0 1 2 3 lo g a m p lit u d e 1 2 5 10 20 50 frequency (hz) meteroid thunder earthquake fig. 6. frequency spectra of meteoroid, thunder and earthquake waves. a time window of 2 s around the signal on the up/down z-sensor from station1 was used. the amplitude is uncorrected. the nyquist frequency is 50 hz. geological survey of denmark and greenland bulletin 33, 2015, 33-36 33© 2015 geus. geological survey of denmark and greenland bulletin 33, 33–36. open access: www.geus.dk/publications/bull relationship between groundwater chemistry and the precambrian basement rocks on eastern bornholm, denmark peter gravesen, rasmus jakobsen and bertel nilsson bornholm is situated south of sweden, in the sorgenfrei– tornquist zone (fig. 1). th e precambrian basement on northern and eastern bornholm consists of diff erent types of granitic and gneissic rocks with pegmatites, aplites and dolerite dykes (callisen 1934). th e age of the granite and gneiss is c. 1455 ma (waight et al. 2012). th is study deals with the østerlars–svaneke area north of paradisbakkerne and focuses on the geology and groundwater chemistry of the groundwater aquifers in the bornholm gneiss, paradisbakke migmatite and svaneke granite (fig. 2). th e geology and groundwater conditions in part of the study area were described by gravesen et al. (2011, 2013, 2014). miljøcenter roskilde (2009) described the groundwater conditions at the østerlars and østermarie waterworks. th e aim of this paper is to show the relations between rock composition, mineral alteration and groundwater chemistry in the low-permeability rocks using existing data from outcrops and boreholes. methods data on geology and groundwater chemistry were obtained from the jupiter database at the geological survey of denmark and greenland. information from more than 200 boreholes reaching basement rocks is included; in addition some data on the quaternary sediments were available. of the 200 boreholes, data on groundwater chemistry from 29 boreholes were available, in some boreholes as a series of analyses. data on fractures were collected from outcrops in the area during fi eld work in 2011. precambrian rock distribution and petrography th ree main rock types are found in the area. th e westernmost type is the medium-grained, foliated bornholm biotite gneiss which is usually dark grey, light grey or red grey; small areas with red grey granite or other colours also occur (callisen 1934; platou 1970). th e dark grey type consists of kfeldspar (35%), quartz (25%), plagioclase (25%), biotite (7%), hornblende (5%) and magnetite (3%) and minor amounts of apatite and traces of zircon, allanite, epidote and calcite (micheelsen 1961). light grey quartz gneiss and quartzitic types have a larger content of quartz and are oft en banded. large crystals of apatite and plagioclase occur. th e gneiss also contains skarn bodies with garnet and epidote or lenses with wollastonite, epidote and garnet. fineto medium-grained paradisbakke migmatite occurs in a restricted area between the bornholm gneiss and the svaneke granite (fig. 1) and consists of almost parallel, light grey granitic quartz-feldspar veins in a darker matrix. th e migmatite consists of k-feldspar (35%), quartz (23%), plagioclase (25%), hornblende (8%), biotite (7%), magnetite (1%) and titanite (1%), zircon and traces of allanite; calcite and epidote occur in the darker part but are rare in the lighter part (micheelsen 1961). th e svaneke granite consists of k-feldspar (36%), plagioclase (26%) and quartz (25%) with biotite (7%), hornblende (2%) and contains minor amounts of magnetite and titanite, palaeozoic sandstone dolerite dyke svaneke granite paradisbakke migmatite bornholm gneiss fault 2 km svaneke øster lars paradisbakkerne 55°6´n 15°2´e l isted nexø østermarie årsda le grisby præstebo quarr y s–t zone bornholm fig. 1. geological map of part of eastern bornholm, modified from varv (1977). inset: bornholm’s location in the sorgenfri–tornquist zone (s–t zone). 3434 as well as traces of apatite, zircon, epidote, allanite, calcite and fl ourite (micheelsen 1961). many small inclusions rich in mafi c minerals also occur and locally the content of apatite and fl ourite in the rock appears to be higher. th e granite is usually coarse-grained and greyish red but occasionally medium to coarse grained and yellow, yellow red or dark grey. th e svaneke granite can be divided into four types and a border facies bordering the gneiss and migmatite (platou 1968, 1970). th e border facies is commonly strongly lineated but also has non-lineated parts and a varying content of dark minerals. svaneke granite type i is light yellow and varies in grain size, and biotite is weakly altered to chlorite; this type occurs around svaneke and listed. svaneke granite type ii covers an area from the north coast and inland between the border facies and type i and types iii and iv. th e rock is medium to coarse grained and among the dark minerals hornblende, biotite and sphene are dominant. th e plagioclase and the dark minerals are weakly altered to unaltered. svaneke granite type iii covers a large area along the east coast from grisby to the palaeozoic sandstone north of nexø and is exposed at several places. th e rock is light yellow to light yellow red. th e svaneke granite type iv occurs as lenses in type iii; seen for example at årsdale. it is a dark red, coarse-grained rock and both the plagioclase and the dark minerals are strongly altered. in types iii and iv the hornblende, pyroxene and biotite may be altered to green chlorite, and magnetite is altered to hematite. especially in the coastal area around årsdale and southwards towards nexø the granite is strongly weathered and altered and contains 3–7% chlorite. th e rock readily disintegrates to form coarse gravel (årsdale gravel) and the fracture surfaces of the granite are covered by clayey material of green chlorite and yellow-brown limonite. fracture systems four fracture systems have been identifi ed in the area. th e fracture systems are seen in quarries in the paradisbakke migmatite along the northern rim of paradisbakkerne, in coastal exposures of the bornholm gneiss towards the north and in exposures of the svaneke granite towards the east and the north (von bubnoff 1942). two vertical fracture sets with orientations nne–ssw and ese–wnw and two horizontal fracture sets are found. th e 3d fracture network of crossing vertical, subvertical, horizontal and subhorizontal fractures forms the groundwater aquifers in the rocks. th e vertical fractures are mainly transport paths for infi ltrating water to the groundwater table, whereas the horizontal fractures form conduits for groundwater over long horizontal distances. horizontal fractures are present at least up to 90 m depth below the ground surface; two other shallower sets of horizontal fractures are also found (gravesen et al. 2014). information from boreholes shows that groundwaa b c d fig. 2. the different basement rock types. a: paradisbakke migmatite with pegmatite body, præstebo quarry. b: banded and folded bornholm gneiss, west of listed, c: svaneke granite, east coast of bornholm, d: weathered and fractured svaneke granite, east of listed. photographs: merete binderup. 35 ter is pumped up from both shallow and deep fractures, the deepest ones 144 m below ground surface. in the weathered rocks iron-bearing minerals are oxidised to yellow-brown clayey iron compounds and the dark minerals are altered to clayey green chlorite that is found on the fracture surfaces. clayey material found on fracture surfaces at depths up to 70 m probably have an impact on the chemistry of the groundwater. groundwater chemistry th e number of chemical analyses of groundwater from wells in the study area is unfortunately rather small and some parameters, especially from the svaneke granite, appear imprecise, e.g. ph values with only one digit. measurements of al concentrations are usually missing, which makes it impossible to assess the saturation state of al-silicates. however, measured concentrations of major cations and measured alkalinity (except for one outlier) appear reliable. speciation and calculations of mineral saturations using the soft ware phreeqc (parkhurst & appelo 2012) based on the parameters available, with reservations for the approximate ph values indicates that groundwater in all samples is slightly supersaturated with respect to calcite. a few analyses included measurements of sulphide and trace metal concentrations and the phreeqc calculations indicate that concentrations of trace metals such as pb and zn are likely controlled by sulphide phases. based on the major cations dissolved in the groundwater, it appears that weathering of the diff erent granite types leads to diff erent water chemistry. in a ternary plot (fig. 3) showing the relation between (charge equivalents) of ca, mg and na+k in the water, the bornholm gneiss is barely distinguishable from samples that come from aquifers in quaternary sand and gravel, whereas water samples from the paradisbakke migmatite and svaneke granite are enriched in mg and to some extent in na. na enrichment is most pronounced in samples collected close to the coast. high na concentrations are correlated with high cl concentrations, which indicates deposition of sea salts by dry and wet deposition. although also present in seawater, high mg concentrations do not show any correlation with cl concentrations, which indicates that mg comes from weathering reactions. th is is examined in more detail in fig. 4, where mg concentration versus alkalinity is plotted, which is a general indicator for the degree of weathering. as infi ltrating water rich in carbonic acid from the soil zone reacts with the rock, the carbonic acid is used and the alkalinity in the water increases. a sample from the svaneke granite with an unusually low alkalinity of 1.6 meq/l is considered an outlier and was removed from the dataset in fig. 4. it should be noted that the groundwater may have passed through several rock types before reaching the borehole where it was sampled, and there are no reliable groundwater ages so the relative weathering rates based on the water chemistry of the diff erent rocks implies that the bulk water samples from the boreholes represent comparable residence times of the water. th ere are only three paradisbakke migmatite samples; they indicate both low and high rates of weathering. two 14c datings have been made on water from the paradisbakke migmatite, the ages are rather uncertain, but the oldest of these dates is from the borehole with the highest mg concentration and alkalinity. based on the low alkalinity of the water, samples from the quaternary aquifers indicate the lowest rate of weathering, which could be expected. of fig. 4. plot of mg concentration versus alkalinity, which is considered an indicator for the degree of weathering. bornholm gneiss quaternary sand and gravel paradisbakke migmatite svaneke granite (near coast) svaneke granite mg% +k% na ca% 80 7070 80 90 70 70 70 70 80 80 90 bornholm gneiss quaternary sand and gravel paradisbakke migmatite svaneke granite type ii (near coast) svaneke granite types iii + iv (near coast) svaneke granite type ii svaneke granite types iii + iv m g ( m m o l/ l) alkalinity (meq/l) 1.2 1.0 0.8 0.6 0.4 0.2 0.0 4 5 6 7 8 9 10 fig. 3. ternary plot showing the relative distribution (on equivalent basis) of major cations (mg, ca and na+k) in water samples from wells in different granite types. part of the triangle was removed for clarity. 3636 the crystalline rocks, the bornholm gneiss appears to have the lowest weathering rate. two boreholes in the bornholm gneiss have been sampled using two pumps, giving water from diff erent depths (rasmussen et al. 2007). th ese samples show that the lowermost samples have much higher concentrations of fl uoride (1.1 and 1.9 mg/l compared to c. 0.6 and 0.4 mg/l for the mixed samples) and boron (34 and 47 μg/l compared to c. 20 μg/l for the mixed samples), presumably released from either amphiboles or biotite, perhaps apatite in the case of fl uoride. th e higher concentrations at depth indicate a longer residence time, but the phreeqc calculations indicate that the water is close to saturation for apatite so using fl uoride concentrations to quantify residence time could be diffi cult. th ere is no obvious relation between the diff erent types of svaneke granite and the apparent weathering rates based on mg and alkalinity concentrations in the groundwater. still, weathering rates appear to increase from the inland svaneke granite types iii andiv over the near coastal svaneke granite types iii and iv to the svaneke granite types i and ii that show the highest mg concentration and alkalinity, though types i and ii visually appear to be the least weathered. still the water chemistry indicates that the svaneke granites weather faster than the other granite type, this could be related to the many small, mafi c, mineral-rich inclusions found in the svaneke granite or could refl ect a general, primary, textural diff erence that leads to increased weathering rates. conclusions in spite of the small diff erences seen in terms of bulk mineralogy of the granitic and gnessic rocks and the subtypes there is a distinct diff erence in the observed groundwater chemistry. based on the major cations, especially mg, it appears that weathering of the diff erent basement rock types leads to diff erent water chemistry. references callisen, k. 1934: das grundgebirge von bornholm. danmarks geologiske undersøgelse ii. række 50, 266 pp. gravesen, p., binderup, m., nilsson, b. & pedersen, s.a.s. 2011: geological characterisation of potential disposal areas for radioactive waste from risø, denmark. geological survey of denmark and greenland bulletin 23, 21–24. gravesen, p., nilsson, b., binderup, m., larsen, t.b. & pedersen, s.a.s. 2013: geology, seismic activity and groundwater conditions at six potential disposal sites for radioactive waste from risø, denmark. geological survey of danmark and greenland bulletin 28, 13–16. gravesen, p., nilsson, b., rasmussen, p. & pedersen, s.a.s. 2014: borehole logs from the precambrian basement on bornholm, eastern denmark: geology and groundwater fl ow. geological survey of danmark and greenland bulletin 31, 15–18. hansen, m. & poulsen, v. (eds) 1977: geologi på bornholm, 96 pp. varv ekskursionsfører 1, københavn: tidsskrift et varv. micheelsen, h.i. 1961: bornholms grundfj æld. meddelelser fra dansk geologisk forening 14, 308–349. miljøcenter roskilde 2009: sårbarhedsvurdering af grundvandsressourcen på nordbornholm, 103 pp. roskilde: miljøcenter roskilde, miljøministeriet. parkhurst, d.l. & appelo, c.a.j. 2013: description of input and examples for phreeqc version 3 – a computer program for speciation, batch-reaction, one-dimensional transport, and inverse geochemical calculations. u.s. geological survey techniques and methods, book 6, chapter a43, 497 pp., available only at http://pubs.usgs.gov/tm/06/a43. platou, s.w. 1968: on the petrophysical properties of granitic rocks. geologiska föreningens i stockholm förhandlinger 90, 427–433. platou, s.w. 1970: th e svaneke granite complex and the gneisses on east bornholm. bulletin of the geological society of denmark 20, 93–133. rasmussen, p., klitten, k., nielsen, s. & jensen, p. 2007: bornholms regionskommune. logging og vandkemi i vandforsyningsboringer, 2006. danmarks og grønlands geologiske undersøgelse rapport 2007/36, 175 pp. von bubnoff , s. 1942: beiträge zur tektonik des skandinavischen südrandes. 2. die älteren granite bornholms im rahmen der svekofennidischen tektogenese. neues jahrbuch für mineralogie, geologie und paläontologie, beilagen-band 87, 277–396. waight, t., frei, d. & storey, m. 2012: geochronological constraints on granitic magmatism, deformation, cooling and uplift on bornholm, denmark. bulletin of the geological society of denmark 60, 23–46. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark; e-mail: pg@geus.dk geological survey of denmark and greenland bulletin 41, 2018, 43-46 43 during the 2013 field season, siliciclastic and carbonate rocks of the lower palaeozoic sedimentary succession of the franklinian basin in amundsen land, central north greenland, were collected for whole-rock geochemical analysis. these data are evaluated here in an attempt to identify possible hydrothermal signatures related to sediment-hosted zn-pb mineralisation, similar to that found in correlative strata at the large citronen fjord deposit located c. 100 km to the eastnorth-east. in this paper, we use the term sedex in a broad sense to describe stratiform, sediment-hosted deposits that formed either by syngenetic (exhalative) processes or by subsea-floor replacement coeval with sedimentation (e.g. emsbo et al. 2016); the term mississippi valley-type (mvt) is used for non-stratiform zn-pb deposits that formed epigenetically during late diagenesis or tectonism (e.g. leach et al. 2010). regional setting the late precambrian to devonian franklinian basin extends c. 2000 km from the canadian arctic islands to eastern north greenland (higgins et al. 1991). in eastern north greenland, this basin fill overlies the proterozoic independence fjord group and the hagen fjord group, corresponding to the passive continental margin of laurentia. the franklinian basin is characterised by a transition from a deep-water trough, with mainly fine-grained siliciclastic strata, separated from shelf carbonates to the south (fig. 1; higgins et al. 1991). as summarised in kolb et al. (2016), zn-pb mineralisation in the franklinian basin resulted from two different events: early exhalative and/or sub-seafloor replacement in deep-water siliciclastic rocks, and late epigenetic mvt mineralisation in shelf carbonates. the present study concerns the potential for zn-pb mineralisation in the lower ordovician to lower silurian amundsen land group, in amundsen land. in the study area, the amundsen land group comprises black bedded chert and laminated mudstone, commonly siliceous, with subordinate thin-bedded siliceous turbidites and greenish siltstone; locally, thick redeposited chert and limestone conglomerate interbedded with thick calcareous turbidites are present (friderichsen et al. 1982). the chert contains radiolarians (higgins et al. 1991), implying that biogenic silica is responsible for the quartz-rich nature of these rocks, and the siliceous mudstone. approximately 100 km east-north-east of the study area, in northern peary land, correlative siliciclastic rocks host the large undeveloped, sediment-hosted citronen fjord deposit (fig. 1; van der stijl et al. 1998), with reported total resources (measured + indicated + inferred), at a 2.0% zn cut off, of 132 mt with 4.0% zn and 0.4% pb (ironbark zinc base-metal and ree anomalies in lower palaeozoic sedimentary rocks of amundsen land, central north greenland: implications for zn-pb potential diogo rosa, john f. slack and hendrik falck greenland 20°w 82°n 30°w 100 km navarana fjo rd escarpme nt navarana fjo rd escarpme nt peary land johannes v. jensen land a b amundsen land silurian (sandy turbidite) cambro-ordovician (siltstone, mudstone) early cambrian (mudstone, sandstone, conglomerate) basement d ee pw at er de po sit s sh el f de po sit s silurian (carbonates) cambro-ordovician (carbonates, minor mudstone) early cambrian (sandstone, mudstone) franklinian basin proterozoic (sandstone, carbonates, dolerite, basalt) quaternary overburden kap washington volcanic rocks cretaceous–cenozoic (lava, pyroclastic rocks) wandel sea basin carboniferous–cenozoic (fluvial/marine sandstone, carbonates, shale) fig. 1. geology of central north greenland, showing locations of sampled section in amundsen land (a) and of citronen fjord zn-pb deposit (b); modified after escher & pulvertaft (1995). © 2018 geus. geological survey of denmark and greenland bulletin 41, 43–46. open access: www.geus.dk/bulletin http://www.geus.dk/bulletin 4444 2012). in the model of slack et al. (2015), this deposit formed predominantly by exhalative processes. younger epigenetic, carbonate-hosted, mvt zn-pb occurrences, found in the carbonate shelf in southern peary land, are related to the migration of basinal brines expelled by tectonism and/or hydraulic head caused by ellesmerian orogenic uplift during the middle to late devonian (rosa et al. 2016). in amundsen land, no carbonate shelf exists, so this mineralisation style is not expected to be present, although effects of the ellesmerian orogeny are well expressed by open to recumbent folds and local thrust faults. methods all samples were collected along one section across strata of the amundsen land group at wgs84 longitude 35°.3647 w and latitude 82°.9655 e (fig. 1). twenty-two samples of silty limestone, dolomitic mudstone and mudstone were analysed using a variety of methods. all data are from acme analytical laboratories ltd. in vancouver, british columbia (canada), except y and rare-earth elements (ree) that were determined at activation laboratories ltd. in ancaster, ontario (canada). detailed information on methods, standards, and uncertainties are given on the respective web sites (www.acmelab.com; www.actlabs.com). complete analyses of all 22 samples are available in appendix a (online excel file). results several samples have distinctive bulk compositions. for major-element oxides, one of three grey mudstones contains slightly high fe2o3 t (7.83 wt%) relative to average shale (6.75 wt%; appendix iv in krauskopf & bird 1995); this sample also has elevated mno (0.14 wt%) in contrast to the other samples that contain <0.05 wt% mno. the three mudstones have uniformly low total s and organic c (<0.8 wt% and <0.7 wt%, respectively). for metals of economic and exploration interest, one mudstone sample is noteworthy for having slightly anomalous zn (174 ppm), pb (29.6 ppm), ni (75.0 ppm) and as (24.7 ppm) relative to average concentrations in shale (zn = 95 ppm; pb = 20 ppm; ni = 68 ppm; as = 13 ppm; krauskopf & bird 1995, appendix iv). one sample of silty limestone has the highest total s (1.27 wt%) and pb (63.0 ppm) among all 22 analysed samples, the latter concentration being highly anomalous relative to the average of 3.1 ppm pb for unaltered limestone (hartree & veizer 1982). abundances of ree vary greatly from 0.6–2.0 × average post-archaean australian shale (paas; fig. 2). most of the mudstone and all of the carbonate-rich samples (silty limestone, dolomitic limestone, calcareous shale) display relatively flat paas-normalised patterns, which are typical of sedimentary rocks from throughout the geological record (e.g. mclennan 1989). however, one mudstone and both siliceous mudstone samples show slight depletion of light rareearth elements (lree). most of the silty limestone samples display slight enrichment of lree. calculated eu anomalies (eu/eu*), relative to paas, range from 0.90 to 1.51; 20 of 22 samples have positive anomalies, the three highest values (1.41–1.51) occurring in silty limestone. these eu anomalies are not an analytical artifact of ba interference on eu (e.g. slack et al. 2004), because no correlation exists between eu/ eu* and ba. also important is the fact that all samples display small negative ce anomalies (ce/ce*), which relative to paas vary from 0.81 to 0.95; most are true anomalies (i.e., unrelated to anomalous la enrichment), based on a discriminant plot of pr/pr* vs ce/ce* (fig. 3). field of shaley carbonates (n = 5) in hangingwall of citronen fjord deposit footwall of citronen fjord deposit 0.1 1 10 la ce pr nd sm eu gd tb dy ho er tm yb lu sa m pl e/ pa as sa m pl e/ pa as a 0.1 1 10 la ce pr nd sm eu gd tb dy ho er tm yb lu b mudstone siliceous mudstone silty limestone calcareous mudstone dolomitic mudstone fig. 2. plots of rare-earth element concentrations of representative samples of early palaeozoic sedimentary rocks from amundsen land group in amundsen land. a: mudstone and siliceous mudstone. b: calcareous mudstone, dolomitic mudstone, and silty limestone. field of samples hosting the citronen fjord deposit are included (slack et al. 2015), for comparison; note that small positive eu anomalies for these samples (1.16-1.29) are not evident due to overlapping patterns. normalisations are to average postarchean australian shale (paas); data from taylor & mclennan (1985). http://www.acmelab.com http://www.actlabs.com http://www.geus.dk/media/19170/nr41_p43-46-appendix-a-whole-rock-analyses.ods 45 discussion the presence in one mudstone sample of slightly high fe2o3 t, zn, pb, ni and as is permissive evidence of a hydrothermal component being present in the basin. the small lree depletion in this sample and in the two siliceous mudstone samples (fig. 2a), likely reflects the dissolution of detrital apatite, which in low-temperature sedimentary environments occurs by interaction with acidic fluids and not typical seawater-derived pore fluids (see slack et al. 2017). the geochemical data for this mudstone sample, namely elevated mno together with very low mo, record sedimentation and early diagenesis in oxic bottom waters (e.g. slack et al. 2017). oxic bottom waters are consistent with the presence of small negative ce anomalies in this sample, in both siliceous mudstone samples, and in most of the carbonate-rich rocks. these conditions, as well as the apparently low availability of h2s in pore fluids beneath the palaeo-sea floor (total s <0.8 wt%), were also proposed by slack et al. (2015) for the host sedimentary rocks during initial formation of the citronen fjord deposit. however, according to their model for that deposit, only after emplacement of debris flows that physically restricted the local basin and sealed off communication with the larger oxic ocean, did the venting of hydrothermal fluids turn the bottom waters anoxic and possibly locally very reducing (euxinic) and allow for sulphide preservation. if this model for the redox evolution of the citronen fjord deposit is correct, an analogous scenario for amundsen land (this study) hinges on verifying the local presence of anoxic to euxinic bottom waters, a requirement as yet unachieved, without supporting evidence from additional sampling and analyses. the presence of small positive eu anomalies in most samples is consistent with a hydrothermal component (e.g. lottermoser 1992). however, other non-hydrothermal processes can also create small positive eu anomalies in sedimentary rocks, both siliciclastic and carbonate. for example, in organic-rich black shales, small positive eu anomalies may form diagenetically in euxinic pore fluids (slack et al. 2017, and references therein), but no evidence of such fluids exists in the geochemically anomalous mudstone, based on its elevated mno (0.14 wt%) coupled with low organic c (0.37 wt%) and very low mo (1.75 ppm) contents, which together indicate oxic (not anoxic or euxinic) bottom waters and pore fluids (see slack et al. 2017). furthermore, toc values lack any correlation with metal concentrations. the relatively high fe2o3 t content of this mudstone sample could be a hydrothermal signature, but might also reflect a detrital component derived from a fe-rich source area. regarding the positive eu anomalies present in all of the carbonate samples, a possible non-hydrothermal origin for this anomaly may be related to a large clay component (tostevin et al. 2016), but this explanation is ruled out by the fact that the samples with the highest eu/eu* values (1.41–1.51) have uniformly low al2o3 (0.49–0.62 wt%). given these observations, we conclude that the small positive eu anomalies reflect a hydrothermal signature, involving the passage of reduced fluids that preferentially carried eu2+ (bau 1991). importantly, a hydrothermal origin has also been proposed by several workers for positive eu anomalies in the carbonate gangue and carbonate-rich wall rocks and country rocks of several stratiform sedex deposits (e.g. slack et al. 2004; frimmel 2009). the inferred hydrothermal component in the early palaeozoic siliciclastic and carbonate rocks of the studied section can be ascribed to either a distal or a proximal source, or both. in the case of a distal source, the likely prolonged (c. 105–106 y) venting of hydrothermal fluids into seawater to form the citronen fjord deposit could account for the eu incorporated into the distal mudstones and carbonates of amundsen land during sedimentation, by mixing of hydrothermally derived eu with seawater. in the latter case, involving a proximal source, the observed base-metal and ree anomalies – both eu and lree – in the samples analysed here could record a hydrothermal signature from a local system of either syngenetic or epigenetic origin. given the apparent lack of organic-rich black shales in the study area with anoxic or euxinic redox signatures, a syngenetic origin for this postulated zn-pb mineralisation is considered unlikely, either by purely exhalative or downward-penetrating brine processes (emsbo et al. 2016; sangster 2018). the occurrence of undiscovered mvt zn-pb deposits is also possible (rosa et al. 2016), but this type of mineralisation is characterised by negative, not 0.90 0.95 1.00 1.05 1.10 1.15 1.20 0.6 0.7 0.8 0.9 1.0 1.1 1.2 c e/ c e* pa as pr/pr* paas true negative ce anomaly true positive ce anomaly fig. 3. plot of ce anomaly (ce/ce*) vs pr anomaly (pr/pr*) for analysed samples of early palaeozoic sedimentary rocks from the amundsen land group in amundsen land. data are normalised to paas. fields after bau & dulski (1996). symbols as in fig. 2. 4646 positive, eu anomalies in carbonate host rocks and gangue minerals (e.g. graf 1984; souissi et al. 2013). in summary, considering all available field and geochemical data, including the lack of evidence for anoxic or euxinic bottom waters during sedimentation, we suggest that the base-metal and ree anomalies highlighted in this study from the amundsen land group, in amundsen land, favour a potential for local sedex zn-pb mineralisation that formed mainly by the sub-sea-floor replacement of carbonate-rich sediments. additional sampling and geochemical analysis are recommended for the study area, in order to better evaluate this mineral potential. acknowledgements this work was financially supported by geus and the ministry of mineral resources of greenland, through the nordzinc project. additional support was provided by the northwest territories geological survey, canada. per kalvig and erik vest sørensen of geus are acknowledged for collaboration during field work. we thank hartwig frimmel (university of würzburg) and steve piercey (memorial university of newfoundland) for helpful reviews. online appendix a: whole-rock analyses of early palaeozoic sedimentary rocks from the amundsen land group in amundsen land. references bau, m. 1991: rare-earth element mobility during hydrothermal and metamorphic fluid-rock interaction and the significance of the oxidation state of europium. chemical geology 93, 219–230. bau, m. & dulski, p. 1996: distribution of yttrium and rare-earth elements in the penge and kuruman iron-formations, transvaal supergroup, south africa. precambrian research 79, 37–55. emsbo, p., seal, r.r., breit, g.n., diehl, s.f. & shah, a.k. 2016: sedimentary exhalative (sedex) zinc-lead-silver deposit model. u.s. geological survey scientific investigations report 2010–5070–n, 57 pp. escher, j.c. & pulvertaft, t.c.r. 1995: geological map of greenland, 1:2 500 000. copenhagen: geological survey of greenland. friderichsen, j.d., higgins, a.k., hurst, j.m., pedersen, s.a.s., soper, n.j. & surlyk, f. 1982; lithostratigraphic frameworks of the upper proterozoic and lower palaeozoic deep water clastic deposits of north greenland. rapport grønlands geologiske undersøgelse 107, 19 pp. frimmel, h.e. 2009: trace element distribution in neoproterozoic carbonates as palaeoenvironmental indicator. chemical geology 258, 338–353. graf, j.l., jr. 1984: effects of mississippi valley-type mineralization on ree patterns of carbonate rocks and minerals, viburnum trend, southeast missouri. journal of geology 92, 307–324. hartree, r. & veizer, j. 1982: lead and zinc distribution in carbonate rocks. chemical geology 37, 351–365. higgins, a.k., ineson, j.r., peel, j.s., surlyk, f. & sønderholm, m. 1991: lower palaeozoic franklinian basin of north greenland. bulletin grønlands geologiske undersøgelse 160, 71–139. ironbark zinc ltd. 2012: citronen january 2012 resource. http://ironbark.gl/projects/greenland/citronen/ kolb j., keiding j.k., steenfelt a., secher k., keulen n., rosa, d. & stensgaard, b.m. 2016: metallogeny of greenland. ore geology reviews 78, 493–555. krauskopf, k.b. & bird, d.k. 1995: introduction to geochemistry, 647 pp., third edition. new york: mcgraw-hill, inc. leach, d.l., taylor, r.d., fey, d.l., diehl, s.f., saltus, r.w. 2010: a deposit model for mississippi valley-type lead-zinc ores. u.s. geological survey scientific investigations report 2010–5070–a, 52 pp. lottermoser, b.g. 1992: rare earth elements and hydrothermal ore formation processes. ore geology reviews 7, 25–41. mclennan, s.m. 1989: rare earth elements in sedimentary rocks: influence of provenance and sedimentary processes. reviews in mineralogy 21, 169–200. rosa, r., schneider, j. & chiaradia, m. 2016: timing and metal sources for carbonate-hosted zn-pb mineralization in the franklinian basin (north greenland): constraints from rb-sr and pb isotopes. ore geology reviews 79, 392–407. sangster, d.f. 2018: toward an integrated genetic model for vent-distal sedex deposits. mineralium deposita, 53, 509–527. slack, j.f., kelley, k.d., anderson, v.m., clark, j.l. & ayuso, r.a. 2004: multistage hydrothermal silicification and fe-tl-as-sb-ge-ree enrichment in the red dog zn-pb-ag district, northern alaska: geochemistry, origin, and exploration applications. economic geology 99, 1481–1508. slack, j.f., rosa, d. & falck, h. 2015: oxic to anoxic transition in bottom waters during formation of the citronen fjord sediment-hosted zn-pb deposit, north greenland. in: andré-mayer, a.-s. et al. (eds): mineral resources in a sustainable world: proceedings of 13th biennial sga meeting (nancy, france), 5, 2013–2016. slack, j.f., falck, h., kelley, k.d. & xue, g.g. 2017: geochemistry of host rocks in the howards pass district, yukon-northwest territories, canada: sedimentary environments of zn-pb and phosphate mineralization. mineralium deposita 52, 565–593. souissi, f., jemmali, n., souissi, r. & dandurand, j.l. 2013: ree and isotope (sr, s, and pb) geochemistry to constrain the genesis and timing of the f-(ba-pb-zn) ores of the zaghouan district (ne tunisia). ore geology reviews 55, 1–12. taylor, s.r. & mclennan, s.m. 1985: the continental crust: its composition and evolution, 312 pp. oxford, uk: blackwell scientific publications. tostevin, r., shields, g.a., tarbuck, g.m., he, t., clarkson, m.o. & wood, r.a. 2016: effective use of cerium anomalies as a redox proxy in carbonate-dominated marine settings. chemical geology 438, 146–162. van der stijl, f.w. & mosher, g.z. 1998: the citronen fjord massive sulphide deposit, peary land, north greenland: discovery, stratigraphy, mineralization and structural setting. geology of greenland survey bulletin 179, 40 pp. authors’ addresses d.r., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k denmark. e-mail: dro@geus.dk. j.f.s., u.s. geological survey (emeritus), national center, ms 954, reston, va 20192 usa. h.f., northwest territories geoscience survey, p.o. box 1320, yellowknife, nwt x1a 2l9 canada. http://www.geus.dk/media/19170/nr41_p43-46-appendix-a-whole-rock-analyses.ods http://www.geus.dk/media/19170/nr41_p43-46-appendix-a-whole-rock-analyses.ods http://ironbark.gl/projects/greenland/citronen/ http://ironbark.gl/projects/greenland/citronen/ mailto:dro@geus.dk geological survey of denmark and greenland bulletin 1, 657-722 east greenland previous page: upper jurassic – lower cretaceous sediments of the kap leslie and hartz fjeld formations in milne land, east greenland (viewed towards the east) – see larsen et al. (2003, this volume) and surlyk (2003, this volume). photo: michael larsen. 659 the jurassic of east greenland: a sedimentary record of thermal subsidence, onset and culmination of rifting finn surlyk the late palaeozoic – mesozoic extensional basin complex of east greenland contains a record of deposition during a period of rhaetian – early bajocian thermal subsidence, the onset of rifting in the late bajocian, its growth during the bathonian–kimmeridgian, culmination of rifting in the volgian – early ryazanian, and waning in the late ryazanian – hauterivian. the area was centred over a palaeolatitude of about 45°n in the rhaetian and drifted northwards to about 50°n in the hauterivian. a major climate change from arid to humid subtropical conditions took place at the norian–rhaetian transition. deposition was in addition governed by a long-term sea-level rise with highstands in the toarcian–aalenian, latest callovian and kimmeridgian, and lowstands in the latest bajocian – earliest bathonian, middle oxfordian and volgian. the rhaetian – lower bajocian succession is considered the upper part of a megasequence, termed j1, with its base in the upper lower triassic, whereas the upper bajocian – hauterivian succession forms a complete, syn-rift megasequence, termed j2. the southern part of the basin complex in jameson land contains a relatively complete rhaetian–ryazanian succession and underwent only minor tilting during middle jurassic – earliest cretaceous rifting. rhaetian – lower jurassic deposits are absent north of jameson land and this region was fragmented into strongly tilted fault blocks during the protracted rift event. the syn-rift successions of the two areas accordingly show different long-term trends in sedimentary facies. in the southern area, the j2 syn-rift megasequence forms a symmetrical regressive–transgressive–regressive cycle, whereas the j2 megasequence in the northern area shows an asymmetrical, stepwise deepening trend. a total of eight tectonostratigraphic sequences are recognised in the rhaetian–hauterivian interval. they reflect major changes in basin configuration, drainage systems, sediment transport and distribution patterns, and in facies and depositional environments. the sequences are bounded by regional unconformities or flooding surfaces and have average durations in the order of 10 ma. they are subdivided into conventional unconformity-bounded depositional sequences with durations ranging from tens of thousands of years, in the milankovitch frequency band, up to several million years. deposition was alluvial and lacustrine in the rhaetian–sinemurian, but almost exclusively marine during the pliensbachian–hauterivian time interval when a marine strait, up to 500 km wide and more than 2000 km long, developed between greenland and norway, connecting the arctic sea and the north sea. coal-bearing fluvial and paralic deposits occur, however, at the base of the onlapping middle jurassic succession in the central and northern part of the basin complex. the sedimentary development is similar to that in the northern north sea and on the norwegian shelf, and east greenland offers important onshore analogues for virtually all of the types of deeply buried jurassic depositional systems of these areas and especially their hydrocarbon reservoirs. keywords: east greenland, jurassic, sequence stratigraphy, pre-rift megasequence, syn-rift megasequence, basin evolution, regional correlation/comparison, onshore moray firth geological institute, university of copenhagen, geocenter copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: finns@geo.geol.ku.dk geological survey of denmark and greenland bulletin 1, 659–722 (2003) © geus, 2003 the jurassic system of east greenland has a long history of exploration and is one of the most well-exposed ancient extensional basin successions known. in jurassic time, an up to 500 km wide and more than 2000 km long epicontinental seaway existed between greenland and norway (fig. 1; ziegler 1988; doré 1992). the deposits of the western margin of the seaway are today widely exposed in east greenland through cenozoic uplift of 1–2 km (christiansen et al. 1992a; hansen 2000; johnson & gallagher 2000; mathiesen et al. 2000). the deposits of the central and eastern parts of the jurassic seaway are deeply buried in the outer shelves of offshore east greenland and norway, with the exception of a small outlier in andøy in northern norway (dalland 1981). in east greenland, the sub-basins form a north–south-trending belt, which is 700 km long, up to 175 km wide at the southern end and narrowing to the north where it passes into the greenland sea (figs 2, 3). the extensive outcrops reveal large-scale geometries of major lithological units, and allow reconstruction of 660 60˚n 55˚n 50˚n pa lae ol at itu de 500 km wollaston forland jameson land mid north sea high moray firthhelm sd ale fa ult in tr arif t h igh middle jurassic land normal fault igneous activity inferred structural high deltaic/shallow marine sandstone offshore marine mudstone ? greenland norway sweden baltic shield laurentian shield ringkøbing–fyn high fig. 1. schematic pre-drift reconstruction of the seaway between greenland and norway showing the position of the mesozoic rift basin of east greenland. based on ziegler (1988) and doré (1992). the main drainage patterns and sediment dispersal systems. detailed sedimentological and sequence stratigraphic analyses can also be performed. the middle– upper jurassic succession has an ammonite record that is unrivalled in the boreal realm, where it serves as a biochronological standard for time correlations (callomon 1959, 1961, 1984, 1993, 1994; surlyk et al. 1973; sykes & surlyk 1976; surlyk 1978a; sykes & callomon 1979; callomon & birkelund 1980, 1982; birkelund et al. 1984; birkelund & callomon 1985). the earlier, first-generation phase of exploration coincided to a large extent with lauge koch’s expeditions, which began in 1926 and ended in 1958. at this time, the main outline of the jurassic succession had become known, most outcrops had been discovered and the rocks dated. these were the foundations laid through numerous expeditions led mainly by lauge koch over about 30 years. the main exception lay in the inland area of the large peninsula of jameson land, which was essentially unknown (fig. 2). excellent reviews of the results obtained during this phase are given by donovan (1957) and haller (1971). in 1968–1971, the jurassic of jameson land and milne land was mapped by geologists from the university of copenhagen under the leadership of the late tove birkelund, and the first coherent lithostratigraphic scheme was erected (surlyk et al. 1973). this was followed by detailed work on the upper jurassic lithostratigraphy, biostratigraphy, palaeoecology and palaeontology of milne land (callomon & birkelund 1980, 1982; fürsich 1982a, b; birkelund et al. 1984; birkelund & callomon 1985). middle jurassic trace fossils and facies of jameson land were studied by heinberg (1970, 1973, 1974), bromley & asgaard (1972) and heinberg & birkelund (1984). rhaetian – lower jurassic palynology was studied by pedersen & lund (1980) and sedimentology by clemmensen (1976) and sykes (1974a, b). a jurassic outlier in liverpool land was described by birkenmajer (1976). middle–upper jurassic syn-rift stratigraphy and sedimentology of the successions of the northern east greenland basin complex were studied by clemmensen & surlyk (1976), surlyk (1977a, b, 1978a, b, 1984, 1989) and surlyk & clemmensen (1983). the combined results of the early exploration and mapping phase and of subsequent more focused studies were synthesised by birkelund & perch-nielsen (1976), surlyk 661 c c' b' a a' 18°w 16°w 74°n 76°n 72°n 22°w 20°w saf pdmf df lle lle cretaceous jurassic triassic permian fault buried deepseated faults stauning alper fault post-devonian main fault dombjerg fault liverpool land escarpment greenland milne land traill ø geographical society ø hold with hope clavering ø th.thomsen land kuhn ø hochstetter forland store koldewey wollaston forland pd m f sa f d f jameson land carl sb er g f jor d hall bredning kong oscar fjord liverpool landhurry inlet aa' b b' c c' 100 km 26°w 22°w24°w28°w 26°w 24°w fig. 2. map of east greenland showing the permian–cretaceous outcrop, major faults thought to have been active during the jurassic and place names mentioned in the text; pre-permian and post-cretaceous geology not indicated. a–a′, b–b′ and c–c′ indicate position of cross-sections on fig. 3. note that the term ‘east greenland’ is used in a broad sense in this paper to include the entire mesozoic rift basin extending from southernmost jameson land in the south to store koldewey in the north. this contrasts with the geus in-house usage in which the northern part of the basin complex, north of kong oscar fjord, falls into ‘north-east greenland’. 662 0 -4 -3 -2 -1 0 1 -3 -2 -1 0 1 10 20 30 40 50 60 70 80 90 km 0 10 20 30 40 50 60 70 80 90 km b b' c c' crystalline basement crystalline basement devonian devonian – lower permian carboniferous – lower permianupper permian – jurassic upper permian – jurassic 0 0 5 10 15 km 1 2 a a' 0 20 40 60 80 100 120 km 0 20 40 60 80 100 120 140 160 km upper bajocian (?) – middle oxfordian upper permian crystalline basement pelion and jakobsstigen formations foldvik creek group km lower volgian – hauterivian upper oxfordian – lower volgian barremian–albian cenozoic plateau basalts hold with hope group wollaston forland group upper permianjurassic triassic bernbjerg formation km km 0 10 20 km 0 2 4 6 twt (sec.) (1978a), surlyk et al. (1981) and surlyk (1990a), and the first low-order sequence stratigraphic interpretation of the complete jurassic succession was presented by surlyk (1990b, 1991a). the most recent phase of investigation of the jurassic of east greenland began in the late 1980s, the main emphasis being on genetic stratigraphy and sedimentology. much of this work was carried out by graduate students of the university of copenhagen and form the subject of ph.d. theses by dam (1991), engkilde (1994), larsen (1995), vosgerau (1997), bjerrum (1999), hansen (1999) and therkelsen (2000). the sequence stratigraphy, sedimentology and ichnology of the uppermost triassic – lower jurassic of jameson land were described by dam (1990a, b), dam & surlyk (1992, 1993, 1995, 1998) and dam et al. (1995), of the middle jurassic of jameson land by engkilde & surlyk (1993, 2003, this volume), of the upper jurassic of milne land by larsen et al. (2003, this volume) and of jameson land by surlyk & noe-nygaard (1991, 1995, 1998, 2000a), surlyk et al. (1993) and larsen & surlyk (2003, this volume). organic geochemistry of various shale-dominated units was presented by dam & christiansen (1990), christiansen et al. (1992b) and krabbe et al. (1994). detailed palynological studies were undertaken by piasecki (1981, 1996), milner & piasecki (1996), koppelhus & dam (2003, this volume), koppelhus & hansen (2003, this volume) and palynological data from the lower–middle jurassic transition of jameson land were included in underhill & partington (1994). jurassic outliers in northern east greenland were examined by stemmerik & piasecki (1990), and the early rift and rift-climax deposits of the wollaston forland area were studied in 1995 by the author in the company of geologists from statoil, saga, amoco, the norwegian petroleum directorate and aarhus university (bojesen-koefoed et al. 1997; petersen et al. 1998; vosgerau et al. 2000; alsgaard et al. 2003, this volume). new outcrops and poorly known successions on traill ø and hold with hope were described by carr (1998), alsen & surlyk (in press), therkelsen & surlyk (in press), vosgerau et al. (in press a, b). possible palaeo-oilfields in the traill ø – geographical society ø area were found and first described by marcussen et al. (1987) and christiansen (1994); they were investigated in more detail by price & whitham (1997) and therkelsen (2000). the present synthesis builds on earlier reviews and studies by donovan (1957), haller (1971), surlyk et al. (1981) and surlyk (1990a, b, 1991a). it incorporates the results of the studies mentioned above, most of which are published in more detail elsewhere in this volume, and further unpublished data of the author and n. noe-nygaard. a marked change from triassic arid continental redbeds to humid drab, dark grey deposits took place in the rhaetian and lasted throughout the jurassic. the succession covered here thus includes the main, upper part of a rhaetian – lower bajocian pre-rift megasequence, and an upper bajocian – hauterivian syn-rift megasequence. a revised lithostratigraphic framework for the jurassic of east greenland is, at the time of printing, in the final stages of preparation by the author and co-workers. the revised and new lithostratigraphic terms shown on figures 4 and 5 and utilised here and in other papers in this book must thus be considered as provisional, pending the publication of the lithostratigraphic revision. separate standard stage names are used for the tethyan and boreal realms for the uppermost jurassic – lowermost cretaceous interval due to strong faunal provincialism among the time-diagnostic guide-fossils. in the tethyan realm, a succession of kimmeridgian, tithonian and berriasian stages are recognised, whereas the kimmeridgian, volgian and ryazanian (or boreal berriasian of russian workers) stages are used in large parts of the boreal realm. the primary standard for the jurassic–cretaceous boundary is conventionally placed at the tithonian–berriasian boundary (casey 1973; birkelund et al. 1984; zeiss 2003, this volume); this definition is followed here. the volgian–ryazanian boundary, which is commonly used as the jurassic– cretaceous boundary in the boreal realm, roughly corresponds to the lower–upper berriasian boundary (zeiss 2003, this volume). thus, when boreal workers talk about the jurassic–cretaceous boundary, they commonly refer to a level in the middle of the berriasian. a slightly smaller problem is presented by the bajocian– bathonian stages in the middle jurassic, where faunal provincialism among ammonites prevents precise correlation between the boreal and other faunal realms. 663 facing page: fig. 3. upper: schematic w–e sections of the mesozoic basin complex of east greenland. note the different tectonic development of the southern and northern areas. a–a´: wollaston forland; b–b´ and c–c´: jameson land. position of sections shown on fig. 2. in jurassic time, jameson land persisted as a broad, gently tilted platform, whereas deep half-grabens were formed in the wollaston forland area during rift culmination. lower: w–e cross-sections showing two different interpretations of the deep structure of the jameson land basin. the mesozoic and deep structure cross-sections are based on vischer (1943), haller (1971), surlyk (1991a) and larsen & marcussen (1992). this problem is discussed further by callomon (2003, this volume). recent strontium isotope work by m. engkilde (personal communication 2000) seems, however, to allow firmer correlation and to corroborate earlier tentative correlations based on ammonites (callomon 1984, 1985). the aims of the paper are to present a genetic stratigraphic interpretation of the upper rhaetian – hauterivian succession of east greenland, to evaluate the factors that governed basin evolution and sedimentation and to compare relative sea-level curves derived for the east greenland basin and the north sea. structural and stratigraphic background a succession of sedimentary basins was formed along the eastern margin of the precambrian shield area of greenland in late palaeozoic – mesozoic times. basin development began in the devonian soon after the culmination of the caledonian orogeny, and caledonian and older structural grains exerted an important control on subsequent basin evolution (surlyk 1977a). the tectonic styles, degree of crustal thinning, and basin evolution vary markedly from south to north (fig. 3; surlyk 1990a). late palaeozoic – mesozoic basin evolution was governed by a system of old ‘first order’ faults some of which were formed by reactivation of caledonian or older thrusts and weakness zones. new faults were formed during continued extensional movements. in the south, a very deep basin was formed in the jameson land area in late devonian – early permian times, recorded in up to 15 km of sediment (surlyk 1991a; larsen & marcussen 1992). this area then acted as a wide, coherent platform in late permian – jurassic times and, for the most part, the succession shows an overall layer-cake geometry without major lateral changes in thickness. subsidence was accommodated mainly along basin margin faults and there was no synsedimentary faulting of the platform. this is in strong contrast to the region north of jameson land, notably the wollaston forland – kuhn ø area, which was fragmented into narrow tilted blocks during middle jurassic – earliest cretaceous rifting (figs 2, 3). the northern and southern areas also differ in another important respect. in jameson land, successive basins are stacked one on top of the other. however, in the areas further north there is a marked eastwards shift in the location of progressively younger basins and depocentres. this trend was first pointed out in the classical but somewhat overlooked study of vischer (1943); a similar, but mirror-image westwards younging of successive basins has recently been recognised on the norwegian shelf (doré et al. 1999). there is thus a marked tendency in the northern north atlantic region for a symmetrical lateral offset of successive axes of crustal thinning towards the eventual line of palaeogene break-up. the east greenland rift basin complex was centred over 45°n in the rhaetian and drifted northwards to reach 50°n in the hauterivian (smith et al. 1994). a major climate change from arid to humid subtropical conditions took place at the norian–rhaetian transition. the different structural style and degree of late palaeozoic crustal thinning along the length of the east greenland basin exerted a strong control on jurassic genetic stratigraphic development. the jameson land platform contains an aggradational rhaetian – lower bajocian pre-rift package overlain by an overall regressive–transgressive–regressive upper bajocian – hauterivian syn-rift succession (fig. 4, facing page 676). this pattern is in contrast to the areas further north, where lower jurassic deposits are absent and a series of backstepping events took place in middle–late jurassic times reflecting increasing rates of extension, culminating in major block-tilting in the middle volgian (surlyk 1991a). the two areas are separated by a system of curved nw–se-trending, deep-seated cross-faults (fig. 2; dam et al. 1995). relative sea-level curves for the two regions are thus roughly similar for the middle and early late jurassic (surlyk 1990b) but deviate for the remaining part of the late jurassic. the jurassic sea level shows a long-term rise with highstands in the toarcian–aalenian, latest callovian and kimmeridgian and lowstands in the latest bajocian – earliest bathonian, middle oxfordian and in the southern area also in the volgian, while the northern area shows continued rise during late jurassic times. tectonic versus eustatic control on relative sea-level changes, and thus on the development of depositional sequences and sequence stratigraphic key surfaces, is a much-debated subject. east greenland provides a good example of how longterm trends in sea level are influenced by tectonic style and intensity of rifting. jurassic stratigraphy of east greenland sedimentary successions are commonly described within the framework of conventional lithostratigraphic schemes. focus on lithostratigraphy by itself, however, may fail to highlight genetic relationships between rock units and hence to develop a thorough understanding 664 of the factors governing basin evolution (surlyk 1991a). a modern lithostratigraphic scheme is much more than a representation of simple mapping units. it should reflect a thorough understanding of the genetic stratigraphy including sequence stratigraphy of the basin fill. it should only be established when the basin and its fill has received detailed study. premature designation of lithostratigraphic units following initial mapping commonly severely hampers succeeding work (see also miall 1990, p. 402). the extensive nature of the jurassic outcrops in east greenland has, however, made possible the construction of lithostratigraphic schemes that to a large degree also reflect genetic relationships between units (figs 4, 5; surlyk et al. 1973; dam & surlyk 1998). this is because the natural boundaries used in the definition of formations and members commonly coincide with sequence stratigraphic key surfaces rather than purely lithological changes. the individual jurassic formations in east greenland are thus to a large extent genetic units bounded by key surfaces such as unconformities, ravinement and flooding surfaces, rather than merely by arbitrary facies boundaries such as the ‘incoming of the first sandstone’ or the ‘first glauconitic horizon’. the key surfaces in most cases have a distinct lithological expression and can be traced laterally over large areas and are thus excellent lithostratigraphic markers. the main difference between genetic stratigraphy and simple lithostratigraphy in the present case thus lies in the choice of formation boundaries, which are mainly defined by sequence stratigraphic key surfaces. it should be stressed, however, that it is their lithological expression and not their interpreted genesis that forms the basis for the lithostratigraphic definitions. a succession of long-term genetic sedimentary units is recognised in the jurassic succession of east greenland (surlyk 1991a; surlyk & noe-nygaard 2000a). each unit consists of related facies associations representing stacked forestepping, aggradational or backstepping units deposited under roughly similar conditions. the boundaries between the units reflect major changes in tectonic style, basin configuration, drainage patterns or overall depositional environments. the boundaries are unconformities, flooding surfaces and regional facies changes marking discrete basin-fill phases and commonly representing significant hiatuses. the packages are similar in nature to the jurassic sequences recognised in the north sea and labelled j0–j7 by rattey & hayward (1993). they have durations between 5 and 20 ma, averaging 10 ma. however, they are not low-order sequences in a strict exxon/vail sense because they are not consistently unconformity-bound, but are instead genetic tectonostratigraphic units in the sense of surlyk (1991b). they correspond roughly to the main basin-fill phases recognised in earlier syntheses of surlyk et al. (1981) and surlyk (1990a). the tectonostratigraphic units provide the basic framework for the descriptions that follow. they are subdivided into conventional depositional sequences (posamentier & vail 1988) and are treated in more detail in studies of the individual units (dam & surlyk 1998; alsgaard et al. 2003, this volume; engkilde & surlyk 2003, this volume; koppelhus & dam 2003, this volume; koppelhus & hansen 2003, this volume; larsen et al. 2003, this volume; larsen & surlyk 2003, this volume; alsen & surlyk in press; therkelsen & surlyk in press; vosgerau et al. in press a, b). tectonic–stratigraphic evolution during the rhaetian–hauterivian two tectonostratigraphic megasequences, j1 and j2, are recognised in the rhaetian–hauterivian of east greenland. j1 includes the rhaetian – lower bajocian and j2, the upper bajocian – hauterivian succession. the megasequences are subdivided into eight tectonostratigraphic sequences, termed j1.1–1.2 and j2.1–2.6 (fig. 4). j1.1–1.2 and 2.1–2.6 are subdivided into depositional sequences labelled j2.6.1 for the lowest sequence in j2.6 etc. the number system deviates from that of surlyk & noe-nygaard (2000a) in the distinction between the two megasequences but the recognised units are the same. the east greenland and north sea successions show great similarities that are highlighted by the use of similar number systems, but there are also significant differences (partington et al. 1993a, b; rattey & hayward 1993). sequences j1.1 (rhaetian–sinemurian) and j1.2 (pliensbachian – lower bajocian) were deposited during overall tectonic quiescence and regional subsidence and represent late post-rift deposition after rift phases in late permian – earliest triassic and early triassic times. these two tectonostratigraphic sequences are restricted to jameson land, and rhaetian–aalenian deposits are not known from other parts of east greenland (fig. 4). they may have been removed by erosion during late early jurassic uplift of northern east greenland followed by subsidence and middle jurassic onlap (surlyk 1977a, 1978a). the absence of j1.1–1.2 north of jameson land seems, however, to be primary as indicated by the occurrence in northern jameson land of proximal facies of the rhaetian–sinemurian 665 666 uppermost triassic – jurassic of jameson land group formation member group formation member raukelv fynselv salix dal sjællandselv salix dal sjællandselv hareelv olympen fossilbjerget vardekløft pelion sortehat sortehat ostreaelv ostreaelv neill klinter gule horn gule horn rævekløft rævekløft kap stewart kap stewart jameson land jameson land jameson land supergroup scoresby sund hall bredning vardekløft scoresby sund hall bredning vardekløft neill klinter raukelv hareelv olympen fossilbjerget pelion rhætelv primulaelv innakajik straight river fynselv langelandselv katedralen zeus hades athene goniomyakløft parnas ugleelv trefjord bjerg skævdal nathorst fjeld/harris fjeld/ lepidopteriselv astartekløft horsedal albuen elis bjerg surlyk et al. (1973) revised lithostratigraphy (provisional) neill klinter group from dam & surlyk (1998) jurassic of milne land formation member hennigryggenhartz fjeld charcot bugt kap leslie astartedal pernaryggen krebsedal gråkløft cardioceraskløft bays elv aldinger elv kosmocerasdal mudderbugt visdal formation member hennigryggenhartz fjeld charcot bugt kap leslie astartedal pernaryggen krebsedal gråkløft cardioceraskløft bays elv aldinger elv kosmocerasdal mudderbugt visdal supergroup group callomon & birkelund (1980) and birkelund et al. (1984) revised lithostratigraphy (provisional) rauk plateau aucellaelv new revised elevated in rank 667 surlyk (1977b) based on ravn (1911) and koch (1929) stemmerik & piasecki (1990) revised lithostratigraphy (provisional) jurassic of store koldewey and germania land group formation member supergroup group formation jameson land jameson land kløft i vardekløft pelion cf. bernbjerg trækpas hall bredning hall bredning vardekløft bernbjerg pelion payer dal jurassic – lowermost cretaceous of hold with hope, clavering ø, wollaston forland, kuhn ø and hochstetter forland surlyk (1977b, 1978b) revised lithostratigraphy (provisional) group formation member group formation member wollaston forland jameson land jameson land jameson land lindemans bugt bernbjerg bernbjerg bernbjerg bernbjerg vardekløft vardekløft vardekløft niesen rigi laugeites ravine niesen rigi laugeites ravine jakobsstigen pelion pelion pelion muslingebjerg jakobsstigen pelion muslingebjerg supergroup group formation member group formation membersupergroup jameson land wollaston forland hall bredning vardekløft payer dal bastians dal parnas olympen fossilbjerget bristol elv surlyk (1977b) revised lithostratigraphy (provisional) jurassic of traill ø and geographical society ø lindemans bugt palnatokes bjerg rødryggen albrechts bugt falske bugt young sund rødryggen albrechts bugt falskebugt young sund ugpik ravine spath plateau palnatokes bjerg fig. 5. revised stratigraphic schemes for the jurassic of east greenland compared to earlier schemes of surlyk et al. (1973), surlyk (1977b, 1978b), callomon & birkelund (1980), birkelund et al. (1984) and stemmerik & piasecki (1990). stratigraphy of the neill klinter group from dam & surlyk (1998). as discussed in the text, the revised lithostratigraphy should be considered provisional, pending publication; the annotations ‘new, revised, elevated in rank’ refer to this comprehensive lithostratigraphic revision. 668 le pi do pt er is r ha et ia n t ha um at op te ri s h et ta ng ia n delta plain alluvial plain alluvial plain flo ra flo ra rhætelv fm r hæ te lv f m pr im ul ae lv f m in na ka jik f m 0 1 2 3 m 0 1 2 3 m 0 1 2 m 0 1 2 m 0 1 2 m 0 1 2 m silt sand gravel clay silt sand gravel clay silt sand gravel clay silt sand gravel clay silt sand gravel clay silt sand gravel clay silt sand gravel clay silt sand gravel clay delta sheet sandstones open lake 10 m 10 m delta sheet sandstone, strike section rhaetian–sinemurian 100 km n milne land jameson land alluvial/delta plain sandstone/mudstone lacustrine sandstone/mudstone incipient barrier island fig. 6. palaeogeographic map of the rhaetian–sinemurian fluvio-lacustrine kap stewart group. characteristic sections through lacustrine, deltaic and alluvial plain deposits shown. based on dam & surlyk (1993). the accompanying legend is also applicable to figs 9, 10, 12, 13, 16–18, 25–29, 32, 36 and 37. 669 coal claystone mudstone sandstone muddy sandstone pebbly sandstone siderised rip-up mudstone clasts/conglomerate conglomerate, sandstone matrix concretion mudstone clasts carbonaceous material pyritic mudstone sharp/erosional or irregular sharp planar gradational sedimentary logs bed contacts cross-lamination/formset incipient wave ripple lamination wave ripple cross-lamination hummocky and swaley cross-stratification no fabric, random imbricated rootlets plant fragments drifted plant stems/logs bivalves gastropods ammonites belemnites brachiopods crinoids echinoderms conglomerate fabric biota weak moderate bioturbation intense arenicolites isp. chondrites isp. curvolithos multiplex diplocraterion habichi diplocraterion isp. diplocraterion parallelum gyrochorte comosa gyrochorte isp. helminthoida isp. helminthopsis magna monocraterion isp. monocraterion tentaculatum ophiomorpha nodosa phoebichnus trochoides planolites isp. skolithos isp. taenidium serpentinum teichichnus isp. thalassinoides isp. escape burrow unidentified sinuous horizontal burrow trace fossils foreset orientation (crosslamination, cross-bedding) crestline orientation of wave ripples bidirectional current indicators current rose coarsening-upwards trend fining-upwards trend palaeocurrents, miscellaneous parallel lamination planar cross-bedding trough cross-bedding cross-bedding with pebbles along foresets structureless slumping rhythmic lamination lenticular bedding wavy bedding flaser bedding mud partings sedimentary structures heterolithic sediments kap stewart group and the pliensbachian – lower bajocian neill klinter group. the reality of early jurassic domal uplift to the north of jameson land, originally suggested on stratigraphic grounds by surlyk (1977a, 1978a), has recently been confirmed on the basis of apatite fission track thermochronology by johnson & gallagher (2000). in contrast, the jameson land area shows no evidence for early jurassic uplift and cooling (mathiesen et al. 2000). the upper bajocian – volgian of jameson land shows an overall regressive–transgressive–regressive development. the initial progradational phase took place in the late bajocian. it was followed by late bathonian transgression leading to maximum flooding in the middle callovian and the late kimmeridgian – earliest volgian punctuated by progradational interludes in latest callovian – late oxfordian times. maximum late kimmeridgian flooding was succeeded by a major regression in the volgian culminating at the volgian–ryazanian boundary, i.e. at the lower–upper berriasian boundary. deposition of j2.1 (upper bajocian – middle bathonian) reflects the onset of rifting. block faulting with gentle tilting began, the depositional basin was extended far to the north, drainage and sediment transport patterns were fundamentally changed, and enormous volumes of sand were shed into the basin. j2.2 (upper bathonian – upper callovian) records increased rates of extension and block tilting, and progressive drowning of the sandy depositional system. j2.3 (uppermost callovian – middle oxfordian) reflects minor progradational episodes, and j2.4 (upper oxfordian – lower volgian) is characterised by a basin margin progradational interlude along the western basin margin in milne land followed by maximum drowning. j2.5 (middle volgian – upper ryazanian) was characterised by rapid progradation, and sandy depositional systems of j2.6 (upper ryazanian – hauterivian) prograded to the basin axis. the areas further north in the wollaston forland region also show initial late bajocian(?) – bathonian progradation, followed by stepwise middle–late jurassic backstepping with flooding events in the late callovian, late oxfordian and kimmeridgian, culminating in intensive block tilting and formation of deeply submerged half-grabens in the middle volgian. j2.2–2.4 are thus developed as in the jameson land area but increased rifting and block tilting resulted in the formation of progressively deeper water half-grabens with conglomeratic mass flow and turbidite deposition of j2.5–2.6, and the end of the jurassic was characterised by stepwise deepening rather than progradation. j1. rhaetian – lower bajocian pre-rift megasequence kap stewart and neill klinter groups the rhaetian – early bajocian was a time of tectonic quiescence and regional subsidence. sedimentation was restricted to jameson land, which behaved as a single coherent platform (fig. 3). in rhaetian time, the facies changed from continental redbeds to drab sandstones, dark mudstones and thin coals reflecting a change from an arid to a humid climate. j1.1. rhaetian–sinemurian: tectonic quiescence, regional subsidence, fluvial and lacustrine deposition kap stewart group an extensive alluvial–lacustrine complex was formed in the late rhaetian and persisted across the triassic– jurassic boundary into the sinemurian (figs 4, 6, 7; harris 1937; pedersen & lund 1980; dam & christiansen 1990; dam & surlyk 1992, 1993; dam et al. 1995; mcelwain et al. 1999; hesselbo et al. 2002). the deposits are up to about 600 m thick (fig. 8), and are placed in the kap stewart group, which in jameson land conformably overlies red, marly mudstones, grey sandstones and carbonates of the fluvial and lacustrine norian – lower rhaetian ørsted dal member (fleming fjord formation; clemmensen 1980a, b). the boundary is unconformable along the south-eastern basin margin in southern liverpool land, where the kap stewart group, which is here only 20 m thick, oversteps the ørsted dal member and rests on the klitdal member (birkenmajer 1976). along the basin margins, the group consists of coarsegrained alluvial plain deposits (innakajik formation), overlain by finer grained delta plain deposits rich in macroplant fossils (primulaelv formation; figs 5, 6). the entire flora found in the kap stewart group was called the scoresby sund flora (harris 1937 and references therein). gymnosperms (cycadophytes, ginkgophytes and conifers) are dominant, and ferns are also common. the rhaetian lepidopteris flora characterises the innakajik formation and the lower primulaelv formation; the hettangian thaumatopteris flora characterises the upper part of the primulaelv formation (harris 1937). the two floras contain about 100 species each, and have only 5–10 species in common. a similar break occurs in the microflora (pedersen & lund 670 1980). the sharp floral break has been interpreted as reflecting species range truncations at a major hiatus (harris 1937; pedersen & lund 1980). this interpretation is probably not tenable. the two highly diverse and still remarkably different floras and the presence of a transition zone, up to about 10 m thick, is here interpreted to reflect a real floral turnover at or close to the triassic–jurassic boundary. the main facies changes occur in the lower part of the lepidopteris flora interval and the floral break is associated with only a minor change in fluvial style (dam & surlyk 1993). this could be the direct sedimentological result of a floral mass extinction. the reality of the floral turnover is corroborated by a study of stomatal changes between the two floras and the derived fourfold increase in atmospheric pco2 suggesting a 3–4°c ‘greenhouse’ warming across the boundary (mcelwain et al. 1999). the lacustrine deposits of the basin centre consist of alternating black, laminated mudstones and sheet sandstones (rhætelv formation; figs 5, 6, 7). the black 671 sheet sandstone sheet sandstone 20 m sheet sandstone sheet sst sheet sst sb sb fs sb sb sb fs fs dolerite dykes fig. 7. alternating deep lake black shales and shallow lake deltaic shelf sandstones of the kap stewart group in horsedal, northern jameson land. sb, sequence boundary; fs, flooding surface. deep-seated fault 24°w 22°w 50 km < 300 m 300–400 m 400–500 m 500–600 m > 600 m 71°30'n 70°30'n 71°n fig. 8. map of jameson land showing alignment of depocentres of the kap stewart group parallel to deeply buried basement faults. modified from dam et al. (1995). mudstones were deposited mainly during periods of lake-level highstand in relatively deep water under poorly oxygenated conditions. in contrast, the sheet sands were formed by delta progradation during lakelevel lowstand and associated forced regression (fig. 9). the sharp lower boundaries of many of the sheet sandstones are lacustrine regressive surfaces of erosion and form master sequence boundaries (posamentier & allen 1999) or bases of falling stage (plint & nummedal 2000) or forced regressive systems tracts (hunt & tucker 1992, 1995). the upper boundaries are coincident high order sequence boundaries, lacustrine flooding surfaces, and transgressive surfaces of erosion. the deep to shallow lake cycles are thus high-frequency lacustrine sequences. the cyclical alternation of the two facies types shows that the lake experienced numerous high-frequency and high-amplitude changes in water level and shifted from open to closed conditions (dam & surlyk 1992, 1993). about 50 deep lake mudstone – shallow delta sandstone cycles have been recognised. the age control is, however, poor and the duration of the cycles cannot yet be estimated, but it is believed that the cyclicity could be of milankovitch type and be climatically controlled. it is noteworthy that the cycles of the lacustrine kap stewart succession are coeval with the wet–dry cycles of the upper part of the newark supergroup in eastern north america where milankovitch-controlled cyclicity is well-documented (olsen et al. 1996). j1.2. pliensbachian – early bajocian: continued regional subsidence, marine, tidal embayment and restricted offshore deposition neill klinter group the kap stewart lake came to an end at the sinemurian– pliensbachian transition when the jameson land area underwent marine transgression (figs 10, 11; dam & surlyk 1998). deposition seems to have continued unin672 formation of delta sheet sands during lacustrine forced regressions clay silt sand gravel clay silt sand gravel hsttime 5 hsttime 1 fsst+lstsb frse time 2 tse time 3 tst time 4 0 1 2 0 1 a 2 b 3 4 m m open lake open lake tst tse late lst sb fsst frse hst proximal distal a distributary channel b terminal lobe hst highstand systems tract frse forced regressive surface of erosion a b fsst falling stage systems tract sb sequence boundary lst lowstand systems tract tse transgressive surface of erosion tst transgressive systems tract fig. 9. a: development of detached lowstand deltas formed during lake level fall and associated forced regression (based on dam & surlyk 1993 and posamentier & allen 1999). b: sequence stratigraphic interpretation of deep–shallow–deep lake cycles of the kap stewart group (rhætelv formation). sections from dam & surlyk (1993); for legend, see fig. 6. 673 transition zone marine sandstone/ mudstone heterolith sand waves shallow marine sandstone n ei ll k lin te r g ro up waveand stormdominated shoreface tidal channel open lake k ap s te w ar t g ro up k ap s te w ar t g ro up delta plain upper shoreface or surf-zone – deposition from longshore currents r æ ve kl øf t fo rm at io n offshore transition beach foreshore – deposition from swash and foreshore ridges offshore transition g ul e h or n fo rm at io n tidal channel tidal channel silt sand gravel clay silt sand gravel clay silt sand gravel clay 1 m 1 m 1 m g ul e h or n fm , e lis b je rg m b g ul e h or n fm , e lis b je rg m b early pliensbachian flooding surface flooding surface sequence boundary sequence boundary 100 km milne land jameson land n fig. 10. early pliensbachian palaeogeography (rævekløft and correlative basal gule horn formations) of east greenland. characteristic sections through main facies types shown. based on dam & surlyk (1998); for legend, see fig. 6. terrupted in the central part of the basin and the flooding is marked by a coarsening-upwards trend from lacustrine shales to fine-grained sandstones with marine bodyand trace-fossils. pliensbachian marine flooding is recorded throughout the northern north atlantic seaway and coincides with a transgressive peak during the long-term late triassic – early jurassic sea-level rise (hallam 1988; haq et al. 1988; surlyk 1990b). the gradualness of the facies transition suggests that the lacustrine basin was low-lying and separated from the sea only by shallow barriers. a major hiatus comprising all of the sinemurian stage occurs, however, along the south-eastern basin margin caused by erosion and non-deposition during a period of relative sea-level fall close to the sinemurian–pliensbachian boundary. this may reflect a combination of a major eustatic fall and relative uplift and tilting of the liverpool land block to the east as shown by the alignment of depocentres along nw–se-trending border faults at the northeastern end of the jameson land platform. the basin had the same extent as during kap stewart time and was bounded by land to the west, north and east with a marine connection to the south (figs 6, 10). in the basin centre, the marine pliensbachian succession comprises stacked parasequences of waveand storm-dominated offshore transition zone and shoreface tidal channel deposits (elis bjerg member, gule horn formation; fig. 4). along the south-eastern margin, the succession consists of coarse-grained foreshore and shoreface deposits (u. jamesoni – p. davoei chronozones, rævekløft formation), which may be separated by offshore transition zone mudstones representing flooding in the i. ibex chron. early pliensbachian source areas were located along western, northern and eastern basin margins. sediment was distributed by tidal currents in wide dune fields, which passed basinwards into restricted waveand storm-dominated shelf environments. a major landwards shift in facies occurred in the late pliensbachian and a wedge of storm-dominated lower shoreface sediments associated with muddy debrite conglomerates were deposited along the south-eastern basin margin (albuen member, gule horn formation; figs 11, 12). this may reflect a sea-level rise accompanied by erosion of shoreline cliffs. the transgressive development came to an end in the latest pliensbachian, marked by an important basinwards shift in facies rep674 ostreaelv fm gule horn fm rævekløft fm tb s hf as al eb 50 m fig. 11. the neill klinter group at harris fjeld, south-eastern jameson land. note alternating shallow marine, sandstone-dominated, and deeper marine, mudstone-dominated sheets representing individual members of the gule horn and ostreaelv formations (pliensbachian–toarcian). as, astartekløft mb; al, albuen mb; eb, elis bjerg mb; hf, harris fjeld mb; s, skævdal mb; tb, trefjord bjerg mb. resented by subtidal channel and subtidal sand sheet deposits (astartekløft member). a remarkable lateral facies change took place in the northern part of the basin where stacked lagoonal parasequences capped by rootlet beds were formed (horsedal member, gule horn formation; fig. 4). the lagoonal deposits covered more than 3000–4000 km2 (fig. 12). seismic data suggest that the position of the barrier protecting the lagoon was controlled by deepseated nw-trending basement faults and the barrier was apparently located over the subtle footwall crest of the main fault (figs 2, 8; dam et al. 1995). 675 late pliensbachian silt sand gravel silt sand gravel silt sand gravel n clay clay clay stacked waveand storm-dominated lagoonal deposits stacked waveand storm-dominated lagoonal deposits 1 m 1 m 1 m subtidal sand sheet h or se da l m em be r h or se da l m em be r a st ar te kl øf t m em be r flooding surface sand waves tidal deltas transition zone marine sandstone/ mudstone heterolith shallow marine sandstone lagoonal mudstone lagoonal sandstone 100 km milne land jameson land traill ø fig. 12. late pliensbachian palaeogeography (horsedal and astartekløft members, lower ostreaelv formation). characteristic sections through main facies types shown. based on dam & surlyk (1998); for legend, see fig. 6. 676 flooding surface silt sand gravel silt sand gravel clay 1 m 1 m n at ho rs t fj el d m em be r waveand stormdominated shoreface flooding surface bioturbated shoreface clay offshore transition zone sk æ vd al m em be r terminal lobe h ar ri s fj el d m em be r early toarcian transition zone marine sandstone/ mudstone heterolith locally mud-rich sand waves shallow marine sandstone 100 km milne land jameson land traill ø n fig. 13. early toarcian palaeogeography (harris fjeld, nathorst fjeld and skævdal members, ostreaelv formation). characteristic sections through the main facies types shown. based on dam & surlyk (1998); for legend, see fig. 6. 677 sb trefjord bjerg mb skævdal mb fig. 14. tidal sandwave sandstones of the trefjord bjerg member (ostreaelv formation). trefjord bjerg, eastern central jameson land. note well-developed sequence boundary (sb) where the shallow marine sandstones overlie deeper marine mudstones of the skævdal member (ostreaelv formation) marked by an erosional unconformity and a seawards shift in facies. sb? fs fs 2 m sortehat fm ostreaelv fm (trefjord bjerg mb) dolerite sills fig. 15. sharp flooding surface (fs) where offshore mudstones of the aalenian – lower bajocian sortehat formation overlie toarcian shallow marine sandstones of the ostreaelv formation (trefjord bjerg member). cryptic sequence boundary (sb) is situated some metres below the top of the ostreaelv formation. neill klinter, south-east jameson land. a new transgressive phase took place in the early toarcian (fig. 13). the tidal channels and lagoonal systems of the previous phase were drowned and offshore transition zone conditions were developed especially in the southern part of the basin. this was followed by renewed progradation of sands deposited in ebb-tidal deltas (harris fjeld member), shoreface (nathorst fjeld member), and tidal channel complexes (lepidopteriselv member). the latter is separated from the lagoonal deposits of the underlying horsedal member by a marine transgressive surface of erosion. in the late(?) toarcian, a uniform sheet of bioturbated sandy mud and muddy fine-grained sand was deposited throughout the basin (skævdal member; fig. 14). this landwards shift in facies reflects drowning of the coarser grained depositional systems caused by a relative sealevel rise. it was succeeded by a marked facies change at a sharp boundary to sand deposited in marine shoreface, tidal channels and dune fields (trefjord bjerg member; fig. 14). the age relations of the last two units are not yet well-resolved; ammonites are scarce and not all precisely located with respect to the lithostratigraphic units. belemnite data suggest a latest toarcian d. levesquei chron age for the fine-grained skævdal member, whereas ammonites from the base of the member indicate an early toarcian d. tenuicostatum chron age (d. semicelatum subchron). palynomorphs suggest a late toarcian – early aalenian age (koppelhus & dam 2003, this volume). belemnites from the overlying sandy trefjord bjerg member suggest an age no older than the latest toarcian d. levesquei chron and palynomorphs suggest a late toarcian – earliest aalenian age (koppelhus & dam 2003, this volume); the top of the trefjord bjerg member is of latest toarcian – earliest aalenian age according to underhill & partington (1994). palynomorph age indications are corroborated by new sr-isotope data (m. engkilde, personal communication 2000), and the sum of evidence points towards a late toarcian – earliest aalenian age for the two members. recent work by hansen (1999), koppelhus & dam (2003, this volume) and koppelhus & hansen (2003, this volume) on the ostreaelv – sortehat formation boundary has demonstrated the presence of an important sequence boundary unconformity in the top part of the trefjord bjerg sandstones (top ostreaelv formation; fig. 15). underhill & partington (1994) suggested that the socalled ‘mid-cimmerian unconformity’ of the north sea 678 100 km aalenian – early bajocian shoreface mfs rs/fs so rt eh at f m o st re ael v fm so rt eh at f m o st re ael v fm so rt eh at f m o st re ael v fm 1 m 1 m 1 m sand f m c gravel silt clay sand f mc gravel silt clay sand f m c gravel silt clay mfs fs mfs fs shoreface offshore offshore transition zone offshore offshore transition zone offshore marine n milne land jameson land traill ø transition zone marine sandstone/mudstone heterolith offshore marine mudstone fig. 16. aalenian – early bajocian palaeogeography (sortehat formation) with sections through characteristic facies types. sections based on hansen (1999); for legend, see fig. 6. fs, flooding surface; mfs, maximum flooding surface; rs, ravinement surface. was absent in jameson land and that there was a continuous succession over the correlative interval. the new data of hansen (1999) suggest that the sequence boundary unconformity within the upper trefjord bjerg member is broadly coeval with the ‘mid-cimmerian unconformity’; it is important to note, however, that the onset of rifting and the associated change in basin configuration started later in east greenland (mid-bajocian) than in the north sea area, corresponding to the boundary between the sortehat formation and the overlying pelion formation. deposition of the upper sands of the trefjord bjerg member was followed by a major landwards shift in facies represented by a thick coarsening-upwards succession of dark grey to black offshore mudstones and subordinate lower shoreface siltstones and very finegrained sandstones of the sortehat formation (figs 4, 15, 16). the lower boundary of the formation is a prominent northwards younging ravinement surface. a maximum flooding interval is situated close to the base of the formation to the south; it rises to the north reflecting higher sediment influx from northern source areas (hansen 1999). deposition took place in relatively deep water with poor oxygenation at the sea-floor during the early phases as indicated by the sparse macrofauna, low diversity trace fossil assemblage, organic geochemistry and palynomorphs. the lower part of the formation contains salinity tolerant dinocysts and deposition probably took place under brackish-marine conditions (krabbe et al. 1994; hansen 1999; koppelhus & hansen 2003, this volume). accumulation of varying amounts of hydrogen-rich material dominated by the alga botryococcus caused periodic oxygen deficiency (hansen 1999). salinity increased gradually with time, but the low abundance of marine macrofossils seems to reflect dissolution rather than primary absence. ammonites are absent, but palynomorphs suggest an aalenian – (?)early bajocian age for the formation (underhill & partington 1994; hansen 1999; koppelhus & hansen 2003, this volume). j2. late bajocian – hauterivian: main mesozoic rift phase vardekløft, hall bredning and wollaston forland groups the most important rift phase in the mesozoic of east greenland began in the late bajocian, intensified through the bathonian–kimmeridgian, culminated in the middle volgian, and waned in the ryazanian–hauterivian (fig. 4). onset of rifting resulted in the complete reorganisation of basin configuration, drainage patterns and marine transport systems. in early jurassic time, the depositional basin was restricted to jameson land (figs 2, 4). large areas along its western margin in milne land and especially in adjacent areas to the north were now transgressed and became part of the depositional basin for the first time since triassic or older times (figs 2, 4, 17). the transgressive onlap onto crystalline basement is particularly clearly seen on milne land and, to the north, in wollaston forland, kuhn ø, hochstetter forland and store koldewey. jameson land responded to rifting rather differently from the areas further north (figs 3, 4). it behaved as a single structural entity and was not broken up into narrower, strongly tilted blocks. subsidence became asymmetrical through slight tilting, however, with main depocentres in the western and north-western part of the basin and with condensation and unconformities towards the south-east and along the eastern margin of the basin. the platform-type subsidence resulted in almost continuous infilling of the new accommodation space added by the gentle tilting. condensed successions developed in the most distal southern areas, and during the late oxfordian – kimmeridgian sea-level highstand a relatively deep-water basin was formed over southern jameson land. the upper bajocian – upper volgian syn-rift succession thus forms an overall regressive–transgressive–regressive cycle. (note that the late volgian is of earliest cretaceous age and roughly corresponds to the early berriasian.) the nature of the boundary between the jameson land platform and the basins further north is uncertain. surlyk (1978a) suggested that it was developed as a nw-trending crossfault along the present day kong oscar fjord between jameson land and traill ø, following earlier notions of bütler (1957) and donovan (1957). combined study of satellite images, reflection seismic data and lower jurassic facies patterns suggests that the boundary is rather complex and follows a series of deep-seated nw-trending fault splays of probable devonian age in northernmost jameson land (figs 2, 8; dam et al. 1995). rifting began at roughly the same time in the areas north of jameson land. the early phase lasted throughout late bajocian(?) – callovian times with relatively gentle block-tilting. rifting intensified in the oxfordian– kimmeridgian, and the main rift phase took place close to the early–middle volgian transition, accompanied by major block-tilting, formation of high, partly subaerial fault scarps and emergence of footwall crests. rifting waned in the late ryazanian (late late berriasian) and 679 came to an end in the hauterivian, with minor faulting and fragmentation of blocks (vischer 1943; maync 1947, 1949; surlyk 1978b). the middle jurassic – lowermost cretaceous succession in the areas north of jameson land thus shows a stepwise backstepping and deepening represented by the change from shallow marine sandstones, over deeper water outer shelf and slope mudstones to coarse-grained deep-water slope apron and basin-floor fan systems. this is particularly well illustrated by the succession of the wollaston forland area (surlyk 1978b, 1984, 1989; surlyk & clemmensen 1983). renewed episodes of probably rift-related faulting 680 shoreline conglomerate/sandstone shallow marine sandstone transition zone marine sandstone/ mudstone heterolith offshore marine mudstone milne land traill ø geographical society ø maximum transgression clavering ø jameson land wollaston forland 100 km 2 m 2 m sand gravel f silt clay m c sand gravelsiltclay pe lio n fm pe lio n fm late bajocian n hold with hope fig. 17. late bajocian palaeogeography (lower pelion formation and correlatives) with sections through characteristic facies types. note major change in extent and configuration of the depositional basin compared to early jurassic setting. this reflects initiation of rifting along right-stepping, mainly n–s-trending faults. sections based on engkilde & surlyk (2003, this volume); for legend, see fig. 6. occurred in early and early late cretaceous times (surlyk & noe-nygaard 2001a); the jurassic fault blocks were further fragmented in the palaeogene and neogene, and old faults strongly rejuvenated. it is thus rather difficult to identify precisely when the individual faults were active during the long series of rift events, and to estimate the width of the fault blocks at any particular time. the outline of the basin was controlled by n–sand nne–ssw-trending right-stepping en echelon normal faults. this resulted in a northwards tapering of the basin. sediment was introduced mainly at relay or transfer zones at the points of en echelon take-over (surlyk 1977b, 1990a, 1991a; surlyk et al. 1981; surlyk & clemmensen 1983). the overall middle jurassic palaeogeography of the region thus had the configuration of two elongate, fault-controlled embayments with their heads in the clavering ø and hochstetter forland areas. however, the exact position of the embayment heads shifted considerably up and down the axes of the embayments, following changes in relative sea level. the head of the southern embayment thus seems to have shifted from a position in western clavering ø during maximum transgression, as shown by recent finds of jurassic deposits in hold with hope (vosgerau et al. in press a), to a position many tens of kilometres to the south within a relatively short time interval. j2.1. late bajocian – middle bathonian: onset of rifting and marine progradation lower pelion, charcot bugt and fossilbjerget formations rift-induced sedimentation began with fluvial pebbly sandstones of the bristol elv formation (figs 4, 5, 18; therkelsen & surlyk in press). the fluvial deposits reach a thickness of at least 155 m and may represent the fill of an incised river valley system in the traill ø area, which had been uplifted and emergent in early jurassic time (surlyk 1977a, 1978a). fluvial deposition took place under overall base-level rise; peat swamps formed during the last phases of valley filling and are now represented by coaly mudstones. the age of the fluvial succession is not well known, but it shows close lithological resemblance and affinity to the overlying marine sandstones of the pelion formation. the presence of the upper bajocian c. borealis chronozone in the lower pelion formation suggests a slightly older, probably late early bajocian, age for the bristol elv formation. it should be noted, however, that in jameson land the lowest c. borealis occurs 35–40 m above the base of the pelion formation. the age of these pre-borealis beds is not yet known but they undoubtedly represent the distal, marine correlative of the bristol elv formation. in jameson land, the dark mudstones of the aalenian – lower bajocian sortehat formation (top j1.2) are unconformably overlain by shallow marine sandstones of the pelion formation (figs 4, 17, 19). an incised valley has been identified at the base of the pelion formation in northern central jameson land and the boundary is a regional, low order sequence boundary (surlyk 1991a; engkilde 1994; engkilde & surlyk 2003, this volume). the base of the pelion formation is nearly isochronous in the jameson land – traill ø area and can be dated to more or less the late bajocian c. borealis chron. the pelion formation is overlain by and passes southwards into silty, micaceous mudstones of the fossilbjerget formation (figs 4, 17, 20). the boundary between the two formations is highly diachronous, younging to the north from the late bajocian to the middle callovian. it is sharp to the south and becomes more gradational to the north; it is commonly developed as a flooding surface and the formation boundary can be considered a northwards younging succession of shingled flooding or ravinement surfaces. the pelion formation – fossilbjerget formation couplet forms a large-scale genetic unit, which thins markedly in a proximal–distal, north to south direction from about 600 to 150 m, but the main decrease in thickness takes place from central to southernmost jameson land in the area of sandstone pinch-out. the main sediment entrypoint was originally thought to be situated in southern hold with hope (surlyk 1977b; surlyk et al. 1981; surlyk & clemmensen 1983). new finds of middle jurassic deposits in northern hold with hope indicate a northwards extension of the head of the embayment probably to western clavering ø during times of maximum flooding (figs 4, 17; vosgerau et al. in press a). the actual position of the entry point, however, fluctuated considerably in a north–south direction depending on changes in sea level. this is reflected by the pronounced breaks in the succession at hold with hope. the basal sandstones of the pelion formation in this area are of earliest callovian c. apertum chron age, but ammonites of the uppermost bajocian c. pompeckji chronozone occur in the basal cretaceous pebbly sandstone (vosgerau et al. in press a). this shows that the area was flooded during maximum c. pompeckji chron transgression and that pre-callovian pelion formation deposits were removed by erosion during one or more bathonian lowstands. palaeogeographic maps for the late bajocian – callovian time interval thus necessar681 ily present snapshots. the position of the coastline and notably of the head of the embayment shifted back and forth for tens to sometimes more than a hundred kilometres following even minor changes in sea level due to the low inclination of the basin floor. the basic model of surlyk (1977b) and surlyk et al. (1981), involving structurally controlled embayments with the main sediment entry points at transfer zones between right-stepping en echelon faults, is thus considered still to be valid. the base of the early syn-rift unit represented by the sandy pelion formation (and correlatives) youngs northwards from the late bajocian in jameson land and traill ø to the late bathonian/early callovian in hochstetter forland. the onlapping pelion sandstones (and correlatives) rest on progressively older rocks in the same direction. thus they overlie lower bajocian strata in jameson land, triassic strata on trail ø, upper permian strata in southern wollaston forland and crystalline basement rocks in northern wollaston forland, kuhn ø and hochstetter forland. sand and silt were distributed by marine, southwards flowing, tidally enhanced currents and transported as far south as southern jameson land and possibly farther south (fig. 17). the inclination of the n–s-trend682 0 1 2 3 4 5 6 7 8 9 10 11 cl si f m sand c f m pebbles c m m 0 1 2 3 4 5 6 7 8 9 cl si f m sand c f m pebbles c braided river braided river crevasse splay braided river floodplain/lake rootlet horizon floodplain/lake fig. 18. sections through lower bajocian(?) fluviatile pebbly sandstones of the bristol elv formation. based on therkelsen & surlyk (in press); for legend, see fig. 6. 683 sortehat fm pelion fm olympen fm fossilbjerget fm pelion fm ostreaelv fm sortehat fm ostreaelv fm mt pelion fig. 19. the mountain of pelion (1200 m) in northern central jameson land showing thickly developed sandstones of the pelion formation. view towards ese. arrow shows position of section shown on fig. 21. hareelv fm fossilbjerget fm pelion fm (ugleelv mb) fs fig. 20. high-angle clinoform-bedded lowstand wedge sandstones of the ugleelv member, lower pelion formation (upper bajocian). katedralen, south-east jameson land. the wedge is topped by a marked flooding surface (fs) overlain by offshore siltstones of the fossilbjerget formation. the formation boundary youngs northwards from the late bajocian to the early callovian reflecting longterm backstepping and eventual drowning of the shallow marine pelion sandstone system. figure (encircled) for scale. 684 pelion fm sb/tse sb/tse fig. 21. thickly developed monotonous upper bajocian – lower callovian shallow marine sandstones of the pelion formation in northern central jameson land. mount pelion, jameson land (see position on fig. 19) fig. 22. typical simple sequence in the upper pelion formation. the geologist stands on lower bounding surface. degree of bioturbation decreases strongly upwards towards the coincident sequence boundary and transgressive surface of erosion (sb/tse). mount pelion, jameson land. sb/tse fig. 23. lag of well-rounded discoidal quartzite pebbles overlying coincident sequence boundary and transgressive surface of erosion (sb/tse) at the top of simple sequence, as shown in fig. 22. mount pelion, jameson land. 685 ing basin axis was very low and deposition was highly sensitive to even small changes in relative sea level or sediment influx. shoreline progradation or retreat over tens to hundreds of kilometres could thus take place within the resolution of single ammonite faunal horizons (durations down to 100 ka; heinberg & birkelund 1984; callomon 1993, 1994; engkilde & surlyk 2003, this volume). the pelion–fossilbjerget couplet has a uniquely resolved middle jurassic ammonite record (callomon 1993), which allows a detailed sequence stratigraphic breakdown of the thick, rather monotonous sandstone succession (engkilde 1994; engkilde & surlyk 2003, this volume). the pelion formation consists almost exclusively of sandstones and subordinate pebbly sandstones of shoreface and shelf origin (figs 19, 20, 21, 22). any fluvial, beach and foreshore deposits were generally eroded in the marine transgressive shoreface during sea-level rise. the fossilbjerget formation includes the correlative fine-grained offshore transition zone to offshore siltstones and mudstones. the dominant depositional motif of the pelion sandstones is a coarsening-upwards sand-dominated shoreface succession commonly capped by a cemented flooding surface (fig. 22). the lower part is burrow-mottled, while the higher parts have preserved physical sediment structures. the top is normally erosional, commonly with a lag of flat, disc-shaped, well-rounded, quartzite pebbles, and locally a high concentration of ammonites or belemnites (fig. 23). the unit superficially resembles a parasequence bounded by flooding surfaces but the nature of the pebble-lag conglomerates and especially the pebble shapes, indicative of beach swash and backwash, suggest that a high-energy beach bypassed the area during sea-level fall, depositing a relatively thick layer of foreshore pebbly sand. fluvial and beach deposits were reworked during subsequent transgression and the top surface of the coarsening-upwards unit is interpreted as a coincident sequence boundary, marine transgressive surface of erosion, and in some cases a maximum flooding omission surface. the units are accordingly interpreted as simple high-order sequences. a total of 28 high-order sequences with average durations of about 360 ka are recognised and are grouped into eight composite lower order sequences termed p1–8 (p for pelion) with an average duration of 1–2 ma (engkilde & surlyk 2003, this volume). the upper c. 100 m bernbjerg fm olympen fm fossilbjerget fm fs w e pelion fm pelion fm (parnas mb) fossilbjerget fm fig. 24. south coast of eastern traill ø showing pelion, fossilbjerget, olympen and bernbjerg formations deposited on an eastwards tilted fault block. fs, flooding surface. bajocian p1–2 sequences (c. borealis, c. indistinctus, c. pompeckji and lower a. arcticus chronozones) consist predominantly of siltstones and very fine-grained sandstones, and include the most basinally positioned shallow marine sandstones, which were deposited during major sea-level lowstands. they are cut out towards the east by the rocks of the c. greenlandicus chronozone. progradation reached southernmost jameson land where the basal pelion member consists of a coarsening-upwards siltstone–sandstone package (fig. 20). its lower boundary above the mudstones of the sortehat formation is difficult to identify correctly. however, the top sortehat mudstones coarsen upwards into a thin sandstone unit. this is overlain by the basal siltstones of the pelion formation, which coarsen upwards into sandstones capped by a marine flooding surface. during succeeding transgressive and especially highstand time intervals, deposition was restricted to the northern, more proximal parts of the elongate basin. the overlying upper bajocian – middle bathonian p3–4 sequences (upper a. arcticus, a. greenlandicus and lower a. ishmae chronozones) stack aggradationally. the same pattern can be recognised in central traill ø but the succession is thinner probably reflecting proximal bypass during lowstands. in south-eastern traill ø, onset of rifting resulted in eastwards-tilting of a fault block, unusual for east greenland where virtually all jurassic fault blocks dip to the west (fig. 24; donovan 1957; carr 1998; vosgerau et al. in press b). the facies on this block change from proximal to distal towards the east (fig. 26). this suggests that the liverpool land high, which formed the eastern border of the jameson land basin did not extend further north into the offshore area east of traill ø. in the wollaston forland area, the pelion formation may be up to about 500 m thick but is poorly dated and the figure includes the payer dal formation, which may be up to 150 m thick. the monotonous sandstone succession shows few if any well-developed cycles or stacking patterns. sediment influx seems to have been so high that available accommodation space was continuously being filled and the succession can be considered an amalgamated multistorey package. deposition took place under strong tidal influence with dominance of ebb currents towards the south-west (surlyk & clemmensen 1983; alsgaard et al. 2003, this volume). j2.2. late bathonian – middle callovian: early rifting and backstepping of the sandy marine systems upper pelion – charcot bugt – fossilbjerget formations the succeeding three sequences of the pelion formation (p5–7, upper a. ishmae – s. calloviense chronozones) show large-scale late bathonian to early callovian backstepping over several hundred kilometres (fig. 25). a short regressive pulse took place in the early callovian marked by southwards progradation of a sandy wedge to central jameson land (parnas member, c. apertum – c. nordenskjoeldi chronozones), but the sandy pelion system was eventually drowned in the middle callovian and became draped with offshore muds of the fossilbjerget formation, which show condensed levels and hiatuses in the most distal offshore parts in southern jameson land. backstepping was probably governed by increased rates of rifting, with onset of gentle block tilting acting in concert with a late middle jurassic eustatic sea-level rise. a similar sequential development can also be recognised further north in the traill ø and wollaston forland areas, but the preserved sediments are more proximal and sand-dominated, and drowning events are less clearly marked (alsgaard et al. 2003, this volume). the progressive middle jurassic northwards transgressive onlap first reached kuhn ø and hochstetter forland in late bathonian – early callovian times (figs 4, 25). the oldest sediments are fluvial conglomerates and pebbly sandstones of the bastians dal formation, which fill a valley system apparently incised in the crystalline basement or formed by the down-tilted western part of the fault block (fig. 4; alsgaard et al. 2003, this volume). the overlying coal-bearing deposits on kuhn ø and hochstetter forland constitute the upper bathonian? – middle callovian muslingebjerg formation (figs 4, 5; clemmensen & surlyk 1976; petersen et al. 1998). fluvial deposition began during early base-level rise reflecting the ongoing transgression, and the precursor peat swamps of the coal beds were formed when the valley system was almost completely filled. the coal beds in hochstetter forland are interbedded with marine shoreface sandstones and the succession can be subdivided into four high-order sequences each beginning with a thick coal seam resting on a rooted paleosol (fig. 26; petersen et al. 1998). peat accumulation took place during the onset of base-level rise and the coal beds represent the lower part of the transgressive systems tract. the overlying shoreface sandstones form the upper part of the transgressive systems tract and the 686 687 10 m proximal to distal facies changes mud sand milne land traill ø geographical society ø hold with hope clavering ø kuhn ø hochstetter forland store koldewey jameson land wollaston forland late bathonian n 100 km shallow marine sandstone transition zone marine sandstone/ mudstone heterolith offshore marine mudstone localised coal seams palaeocurrents fig. 25. late bathonian palaeogeography (upper pelion and fossilbjerget formations and correlatives) with schematic sections showing the proximal–distal facies development (sections from engkilde & surlyk 2003, this volume). note northwards backstepping of the sandy shallow marine pelion system by comparison with fig. 17. based on surlyk (1977b, 1990a), surlyk et al. (1981) and vosgerau et al. (in press a); for legend, see fig. 6. progradational highstand systems tract. so-called ‘dullingupwards’ cycles in the coal beds represent parasequences or possibly higher order sequences (fig. 26). the final valley-fill stage, including formation of peat swamps, was succeeded by marine flooding in the late callovian and deposition of shoreface and shelf sandstones of the payer dal formation. 688 cl si f m sand c gr 0 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 wave ravinement surface tidal ravinement surface wave ravinement surface flooding surface flooding surface flooding surface m shoreface tidally influenced coastal plain shoreface coastal plain fall rise relative sea level relative sea level relative sea level ds ds ps sq fs sb ps sq fs sb ps sq a a b b fall risefall rise se am 2 se am 2 se am 3 sequence boundary sequence boundary sequence boundary higher order cycles seam 3 seam 2 fig. 26. coal-bearing succession of the upper bathonian(?) – lower callovian muslingebjerg formation at the head of the rifted embayment in hochstetter forland showing a sequence stratigraphic interpretation. on the right, two alternative highresolution sequence stratigraphic interpretations are given of dulling-upwards cycles in coal seam 2: the cycles can be interpreted either as a backstepping parasequence set (a) or as a backstepping composite sequence (b) consisting of simple sequences. sq, sequence; sb, sequence boundary; ps, parasequence; fs, flooding surface; ds, drowning surface. based on petersen et al. (1998); for legend, see fig. 6. farthest to the north, on store koldewey, crystalline basement is onlapped by middle bathonian to lower callovian sandstones of the pelion formation (a. ishmae, a. cranocephaloide, c. apertum and p. koenigi chronozones). the upper part of the succession is a correlative of the spath plateau member of hold with hope (piasecki et al. in press). a spectacular sandy succession is found to the south in milne land (charcot bugt formation) where it onlaps crystalline basement (figs 4, 5; callomon & birkelund 1980; larsen 1995; larsen et al. 2003, this volume). the formation is dominated by thick, high-angle clinoform beds which may represent ebb tidal deltas or possibly shelf-margin deltas or wedges, as they give way distally to, and are overlain by, grey deep-water mudstones of the lower kosmocerasdal member (kap leslie formation) which is the western correlative of the fossilbjerget formation of southern jameson land (fig. 4). the main basin margin fault of northern and central jameson land trends north–south and probably continues offshore into hall bredning between milne land and southern jameson land (fig. 2). the milne land succession thus seems to occupy a position over the footwall forming the western border of the wide jameson land block, which was gradually down-tilted towards the west. if this interpretation is correct, it corroborates the interpretation of the charcot bugt clinoform beds as shelf-margin wedges deposited on a depositional slope below wave base. very similar deposits form a stack of shelf-edge sand bodies in the volgian of southern jameson land (surlyk & noe-nygaard 1991, 1995). backstepping of the pelion formation was probably governed by increasing rates of rifting, block tilting, and deepening of the water over the hanging wall, acting in concert with a late middle jurassic eustatic sealevel rise (hallam 1988; haq et al. 1988). the main depositional motif in the northern areas was similarly a stepwise backstepping of progradational, coarseningupwards sandstone-dominated packages separated by drowning surfaces. j2.3. late callovian – middle oxfordian: renewed marine progradation olympen – upper charcot bugt – jakobsstigen – payer dal formations the large-scale regional drowning of the sandy marine pelion system in the middle callovian was succeeded by progradational episodes in the latest callovian and middle oxfordian (fig. 27). at this time, the basin topography in the jameson land area changed from a ramp to a shelf-break type. the shelf was located over traill ø and northern jameson land, the e–w-trending shelfbreak was situated in northern central jameson land, and the relatively deep-water basin in southern jameson land (fig. 28). the first regressive event was heralded by the incoming of thick massive sands (upper callovian athene member of the olympen formation, p. athleta chronozone; fig. 29). they were triggered from the front of a sandy shelf-margin wedge and transported by sediment gravity flows to the base of slope. the rather uniform massive sandstones belong to an undifferentiated late highstand – falling stage – lowstand systems tract, and a sequence boundary cannot be identified. they are sharply overlain by a thick succession of black mudstones (hades member, olympen formation, lower – lower middle oxfordian, including the q. mariae chronozone), and the boundary is interpreted as a drowning surface, which probably passes updip into the transgressive surface formed by drowning of the shelf-margin wedge (larsen & surlyk 2003, this volume). the hades member mudstones contain a maximum flooding surface of early oxfordian, q. mariae chron age (fig. 29). they are overlain by the second unit of parallel bedded to massive base-of-slope sandstones, probably of late highstand origin, which are sharply overlain by shallow marine delta front sandstones belonging to the falling stage and lowstand systems tract. this upper sandstone unit forms the zeus member of the olympen formation (lower–middle oxfordian, q. mariae – c. densiplicatum chronozones). key surfaces are difficult to identify, and a sequence stratigraphic interpretation is hampered by the deep-water nature of most of the succession. the correlative deposits on traill ø to the north consist of shallow marine sandstones. similar progradational phases occur in milne land at the western basin margin where they are represented by the top of the shallow marine charcot bugt formation (figs 4, 5; callomon & birkelund 1980; larsen et al. 2003, this volume). the late callovian and middle oxfordian regressive pulses can also be recognised in the areas further north where the succession is developed in shallow marine facies. in southern wollaston forland, the tidally influenced high-energy pelion sandstones are sharply overlain by the finer-grained, cyclic and somewhat enigmatic lower–middle oxfordian jakobsstigen formation (figs 28, 30; bojesen-koefoed et al. 1997; vosgerau et al. 2000). the cycles are a few metres thick and consist of 689 690 10 m sand pebbles cl si f m c f c milne land traill ø geographical society ø hold with hope clavering ø kuhn ø hochstetter forland store koldewey 100 km jameson land wollaston forland pe lio n fm fo ss ilb je rg et f m h ar ee lv f m late callovian fs fs fs n shallow marine sandstone transition zone marine sandstone/ mudstone heterolith offshore marine mudstone fig. 27. late callovian palaeogeography (top pelion – fossilbjerget formation level) with section through characteristic facies types. note maximum drowning of the sandy shallow marine pelion system. based on surlyk (1977b, 1990a), surlyk et al. (1981) and vosgerau et al. (in press a); section based on engkilde & surlyk (2003, this volume). for legend, see fig. 6. 691 milne land traill ø geographical society ø hold with hope clavering ø kuhn ø store koldewey 100 km jameson land wollaston forland hochstetter forland middle oxfordian sand f m cmud sand f m cmud mud f m c basin floor o ly m pe n fo rm at io n slope apron a th en e m em be r h ad es m em be r sand 10 m ja ko bs st ig en f m ( lo w er ) ja ko bs st ig en f m ( up pe r) shoreface shallow marine sandstone transition zone marine sandstone/mudstone heterolith offshore marine mudstone deep-water marine sandstone bodies locally mud-rich 5 m 5 m n fig. 28. middle oxfordian palaeogeography (upper charcot bugt, olympen and jakobsstigen formations) with sections through characteristic facies types. note development of a shelf-break topography in jameson land. sections based on surlyk & clemmensen (1983) and larsen & surlyk (2003, this volume); for legend, see fig. 6. thin, sharp-based, dark, organic-rich, laminated mudstones overlain by marine coarsening-upwards, burrowed heteroliths and sandstones. the mudstones do not contain marine microor macrofossils and organic geochemistry also suggests terrestrial deposition. chondrites burrows are conspicuous in the dark mudstones but they are filled with light sand from the overlying shoreface deposits and were thus formed after marine flooding. the organic material of the mudstones includes two populations of organic carbon particles formed by ground fires and high-temperature wild fires. the heteroliths and overlying, commonly cross-bedded, sandstones were deposited under fully marine offshore transition zone and shoreface conditions. the rhythmic alternation between the two facies was caused by relatively high-frequency sea-level fluctua692 10 m fo ss ilb je rg et f m o ly m pe n fm a th en e m b h ad es m b z eu s m b c. densiplicatum chronozone (middle oxfordian) q. mariae chronozone (lower oxfordian) lower–middle p. athleta chronozone (upper callovian) p. athleta chronozone (upper callovian) k. jason chronozone (upper callovian)mud f m c gr sand mud f m c gr sand c. 900 m delta top delta top slope gully basin floor/slope delta front basin floor/slope slope apron basin floor drowning surface fig. 29. sections through the uppermost callovian – middle oxfordian olympen formation showing two progradational events marked by incoming of massive slope sandstones (and delta front and top sandstones in the upper case) separated by a drowning surface of early oxfordian q. mariae chron age. based on larsen & surlyk (2003, this volume); for legend, see fig. 6. 693 tions (vosgerau et al. 2000). the erosional bases of the organic-rich mudstones represent high-order sequence boundaries formed during the lowest stand of sea level. the mudstones were deposited in lakes on a flat, lowlying coastal plain during onset of base-level rise. accelerated base-level rise resulted in marine flooding and termination of coastal plain mud deposition; the overlying marine sandy deposits were deposited during highstand progradation. the organic carbon particles show that groundfires and crownfires were common during early base-level rise, when the area was the site of a lowlying coastal plain with shallow-water lakes. in contrast there are no indications of wildfires during the time of marine highstand deposition. it may thus be speculated that the high-frequency sea-level fluctuations were associated with climatic humid–dry cycles and that the crown fires were associated with the transitional period between the two climatic regimes (bojesen-koefoed et al. 1997; vosgerau et al. 2000). these observations accord well with general circulation models (gcms), which show the area to have very wet winters and very dry summers (sellwood et al. 2000). the wet season would be expected to start with major storms and lightning strikes, starting forest fires. the high-frequency cycles stack into two lower order coarsening-upwards cycles, which may be correlatives of the two progradational cycles of the olympen formation in jameson land. sandstones with pebbly lags of the payer dal formation have recently been identified at hold with hope (vosgerau et al. in press a). they are of late oxfordian a. glosense chron age and overlie pelion formation sandstones of late callovian pre-p. athleta chron age. this suggests the presence of a hiatus, another example of the incomplete nature of the succession at the head of the southern embayment. further north on southern kuhn ø, the upper callovian – middle oxfordian deposits are developed in fully marine facies. the sandstones of the pelion formation are overlain by a coarsening-upwards succession of offshore transition zone heteroliths to shoreface sandstones of the payer dal formation (fig. 31). the age of this unit is not well-known but dinocyst data suggest a middle or late callovian – late oxfordian age. in hochstetter forland, the coal-bearing muslingebjerg formation is directly overlain by the marine sandstones of the payer dal formation, the base of which is dated to the late callovian p. athleta chron, whereas the late oxfordian a. glosense or a. serratum chrons have been demonstrated high in the formation (sykes & surlyk 1976; petersen et al. 1998). bernbjerg fm ugpik ravine mb jakobsstigen fm pelion fm albrechts bugt mb basalt ds ds unc / sb wsw ene fig. 30. the north slope of the valley of cardiocerasdal, wollaston forland, showing stepwise backstepping progradational units of the middle–upper jurassic pelion, jakobsstigen and bernbjerg formations separated by major drowning surfaces (ds). unc/sb, unconformity/sequence boundary. j2.4. late oxfordian – early volgian: increasing rifting and marine flooding hareelv, upper kap leslie and bernbjerg formations the shelf-margin wedges of j2.3 form the youngest preserved deposits in central jameson land (olympen formation; fig. 29). younger deposits of j2.4 exposed in southern jameson land only include slope, base-ofslope and basin-plain deposits (fig. 32). they comprise black mudstones with large bodies of massive sandstones (hareelv formation; figs 32, 33). the sandstones were transported by sediment gravity flows ranging from sandy debris flows to high-density turbidity currents. deposition took place on the slope, base-of-slope and basin floor. some sands were deposited in steep-walled slope gullies, whereas others were loaded into the black slope and basinal muds. the sands were strongly modified by post-burial liquefaction and intrusion into the surrounding mudstones forming a spectacular largescale intrusive complex with extremely irregular sandstone lenses, dykes and sills (figs 33, 34, 35; surlyk 1987; surlyk & noe-nygaard 1998, 2001a, 2001b). remobilisation and intrusion of sand were probably triggered by earthquakes associated with increased rifting. the upper oxfordian part of the succession has roughly equal amounts of mudstone and sandstone, whereas the kimmeridgian part is dominated by massive sandstones. otherwise, the succession is completely chaotic and shows no vertical or lateral trends. key surfaces cannot be identified, and a sequence stratigraphic breakdown has not been possible. a correlative and somewhat shallower water offshore shelf succession is exposed in milne land at the 694 wollaston forland kuhn ø hochstetter forland ns bernbjerg fm payer dal fm pelion fm bastians dal fm caledonian basement permian ugpik ravine mb ? ds/rs ds/rs ds/rs ts muslingebjerg fm offshore marine mudstone transition zone heterolith shallow marine sandstone non-marine (fluvial) sandstone/mudstone coal fig. 31. schematic s–n axial section showing the stepwise backstepping of the middle–upper jurassic succession in the wollaston forland area. progressively, finer-grained progradational units are separated by extensive drowning surfaces. the formations are thus also genetic units. they commonly contain facies in their top part which are similar to those occurring more distally in the underlying unit. this illustrates the difficulties in applying simple lithostratigraphic principles. the whole succession consists proximally of stacked, amalgamated sandstone bodies and distally of clearly demarcated, stacked coarsening-upwards mudstone–heterolith–sandstone units. ts, transgressive surface; ds/rs, drowning surface/ravinement surface. based on alsgaard et al. (2003, this volume). 695 5 m 5 m milne land traill ø geographical society ø hold with hope clavering ø kuhn ø hochstetter forland store koldewey 100 km jameson land wollaston forland sand f mcmud sand f mcmud ds ds lower shoreface – offshore transition zone lower shoreface be rn bj er g fm u gp ik r av in e m b offshore transition zone offshore 5 m 5 m pa ye r d al f m late oxfordian intrusive sands n gully-fill sand sandstone mudstone shallow marine sandstone transition zone marine sandstone/ mudstone heterolith offshore marine mudstone deep-water marine sandstone bodies ? fig. 32. late oxfordian palaeogeography showing sections through characteristic facies types and examples of remobilised sandstones to the right. based on surlyk (1977b, 1990a), surlyk et al. (1981) and surlyk & clemmensen (1983); for legend, see fig. 6. 696 elongate amalg. sst. body n s mst mst mst mst mst sst sst sst sst sandstone body sst & mst fig. 33. large pod-shaped sand body composed of several amalgamated massive deep-water sandstone units. the sands (sst) were liquefied post-burial and intruded into adjacent mudstones (mst). upper oxfordian, hareelv formation, south-east jameson land. sandstone body seen in profile in the northern slope is about 50 m thick. fig. 34. sheet-like massive sandstone representing one sedimentation unit sandwiched between black mudstones. this type of sandstone body is interpreted as having undergone relatively minor post-depositional remobilisation. it corresponds to the sandstone lobe deposit of surlyk (1987). eleven-year-old boy (encircled) for scale at the base of the sandstone. hareelv formation, south-east jameson land. western basin margin (figs 4, 5, 32). the late oxfordian drowning can also be recognised in this area but was rapidly interrupted in mid late oxfordian time by progradation of a shelf sand body (aldinger elv member of the kap leslie formation, a. glosense – a. serratum chronozones) (figs 4, 5; fürsich & heinberg 1983; birkelund et al. 1984). it overlies thick mudstones of the kosmocerasdal member (lower callovian – upper oxfordian, s. calloviense – a. glosense chronozones) to the east in a seawards direction, but wedges out in a landwards direction. the base of the coarsening-upwards sand body is a possible regressive surface of erosion formed during falling sea level. the top is a marine transgressive surface of erosion formed during renewed sea-level rise in the late oxfordian (a. regulare chron). the richly fossiliferous sandstones of the aldinger elv member were originally interpreted as representing an offshore bar (fürsich & heinberg 1983). this interpretation is difficult to uphold because of the basinally-isolated position of the sand body which is encased in offshore mudstone. it is here interpreted as representing a shelf-margin wedge formed during a short interval of rapid sea-level fall followed by rapid transgressive drowning and associated erosion. the aldinger elv member is overlain by a coarsening-upwards succession of siltstones and glauconitic, fine-grained sandstones belonging to the upper oxfordian – lowermost kimmeridgian bays elv member (a. regulare – p. baylei chronozones; figs 4, 5). this is followed by offshore siltstones of the lower kimmeridgian cardioceraskløft member (r. cymodoce – a. mutabilis chronozones) and black, laminated mudstones of the kimmeridgian – lower volgian gråkløft member (a. eudoxus – a. autissiodorensis – p. elegans chronozones) marking maximum late jurassic transgression (fig. 36; birkelund et al. 1984). the upper oxfordian – kimmeridgian succession comprising the aldinger elv, bays elv, cardioceraskløft and gråkløft members shows a strongly backstepping stacking pattern. a similar pattern can be recognised everywhere in east greenland and represents the end of the transgressive part of the middle–late jurassic regressive–transgressive–regressive megacycle in the jameson land – milne land area (fig. 4). maximum flooding persisted to the end of the earliest volgian (p. elegans chron), interrupted by minor regression in the latest kimmeridgian (probably a. autissiodorensis chron). the time interval was characterised by deposition of black, laminated, organic-rich mudstones of the gråkløft member, and represents the highest relative sea level in the jurassic of the jameson land – milne land area (fig. 4). correlative deposits represented by the black mudstones of the upper oxfordian – kimmeridgian bernbjerg formation are poorly preserved in the traill ø region, where they subcrop strongly erosional mid-cretaceous unconformities. late jurassic backstepping continued in the wollaston forland region, and marked drowning surfaces separate the progressively finer-grained and deeper water units (figs 30, 31). the maximum flooding interval comprising the upper kimmeridgian – lowermost volgian a. eudoxus – p. elegans chronozones is thickly developed and consists of black laminated mudstones. sequence stratigraphic interpretation is hampered by the offshore mudstone-dominated nature of the succession. this problem has also been noted in studies of the kimmeridgian of northwest europe (wignall 697 10 cm fig. 35. sandstone dykes and sills in black shale. note cross-cutting relationships indicating several intrusive events. the youngest subvertical dyke is folded due to compaction of the shales. hareelv formation, south-east jameson land. 698 10 m mud sand vf fmc milne land traill ø geographical society ø hold with hope clavering ø kuhn ø hochstetter forland store koldewey 100 km jameson land wollaston forland b er nb je rg f m g rå kl øf t m b c ar di oc er as kl øf t m b turbidites flooding surface 5 m kimmeridgian sand f m cmud transition zone marine sandstone/ mudstone heterolith offshore marine mudstone deep-water marine sandstone bodies n fig. 36. kimmeridgian palaeogeography with sections through characteristic facies types. based on surlyk (1977b, 1990a), surlyk et al. (1981) and surlyk & clemmensen (1983); for legend, see fig. 6. 1991; hallam 1997; taylor et al. 2001). the dominant vertical motif is a coarsening-upwards unit topped by a drowning surface, which may or may not be erosional (figs 31, 32). the cycles thus represent parasequences or distal sequences. farthest to the north, on store koldewey, crystalline basement is onlapped by upper oxfordian – lower kimmeridgian sandstones of the payer dal formation (a. serratum, a. rosenkrantzi and a. mutabilis chronozones) overlain by lower kimmeridgian mudstones of the bernbjerg formation (a. mutabilis chronozone; piasecki et al. in press). j2.5. middle volgian – late ryazanian: rift culmination, rapid progradation in the south, block tilting and gravity flow deposition in the north raukelv and lindemans bugt formations rapid, large-scale progradation characterised the end of the jurassic period in the jameson land area (fig. 37). the succession in milne land at the western basin margin allows a precise dating of the onset of relative sea-level fall to the early volgian p. wheatleyensis chron. the black mudstones representing maximum flooding within the j2.5 sequence (gråkløft member) are sharply overlain by a succession of fine-grained, muddy, coarsening-upwards sandstones of late early – middle volgian age (krebsedal and parnaryggen members; figs 4, 5). the base may represent sea-level fall and associated forced regression. the succession becomes glauconitic at the top, and coarse-grained sandstones occur at several levels. it is capped by a sharp drowning surface of middle volgian c. anguinus chron age. drowning was followed by progradation of coarsegrained, coarsening-upwards sandstone units of middle volgian age, separated by drowning surfaces. in milne land, the sandstone beds are placed in the lower hennigryggen member of the hartz fjeld formation (fig. 38). they are fineto medium-grained, well-sorted, show high-angle clinoform bedding and contain abundant marine trace fossils and large plant fragments. at the top there is a major hiatus between sandstones of the middle volgian l. groenlandicus chronozone (lower hennigryggen member) and lower(?) valanginian, tolliabearing sandstones (upper hennigryggen member). volgian deposits are extremely well-exposed and thickly developed in southern jameson land, where they form a forestepping stack of coarse-grained, pebbly sandstone units constituting the raukelv formation (figs 4, 5, 39–43; surlyk & noe-nygaard 1991, 1995). six major progradational units are recognised in the formation, representing falling stage, lowstand or lowstand wedge systems tracts. they are separated by thin transgressive sandstone sheets. the direction of progradation was mainly lateral, away from the coastline towards the east in contrast to the axial progradation shown by the pelion and olympen formations. the beds show high-angle clinoforms, and are about 15–30 m thick in most of the outcrop area but may reach a thickness of up to 50 m towards the east in a seawards direction. in some cases, it is possible to trace the beds towards the west where they wedge out. the thickest bed shows sigmoidal clinoforms and also has the highest content of marine fossils, notably thick-shelled bivalves and ammonites (fynselv member). the other clinoform beds mainly show tangential clinoforms and marine body fossils are less common. the youngest clinoform bed (rauk plateau member) shows the most proximal character (fig. 39). it is very coarse-grained, lacks body fossils, and the top surface is incised by a network of channels and valleys draped by lags of medium to coarse pebbles and scattered wood fragments. the raukelv formation contains abundant evidence of tidal activity, and the dominant tidal current direction was coast-parallel towards the south with a subordinate current towards the north. the tops of the clinoform beds are draped by lags of well-rounded, discoidal quartzite pebbles with diameters normally in the order of about 1 cm. metre-long hairpin u-burrows of diplocraterion habichi which are characteristic of omission surfaces, descend from the sharp upper surfaces, normally in great density. ammonites, bivalves, belemnites and crinoids are common in beds immediately overlying the omission surfaces. the coarse-grained pebbly clinoform beds of the raukelv formation are interpreted as shelf-margin wedges that prograded seawards below wave base (surlyk & noe-nygaard 1991, 1995). very similar largescale, high-angle clinoform-bedded sand bodies have recently been described from the upper pliocene – lower pleistocene of southern italy; pomar & tropeano (2001) termed these sand bodies ‘transition-slope’ deposits and considered them to represent sediment avalanches swept basinwards from the shoreface and deposited below wave base. the positions of the fluvial entry points of the angular, poorly sorted quartz sands are not known, but were probably somewhere at the western basin margin north of the present location of the shelf-margin clinoform 699 700 5 m n milne land traill ø geographical society ø hold with hope clavering ø kuhn ø hochstetter forland jameson land wollaston forland 100 km ? shallow marine sandstone transition zone marine sandstone/ mudstone heterolith deep-water marine conglomerate/ sandstone sand gravel vfsicl f m c vc vf f m c vc 5 m middle volgian 1 m shelf margin wedge rapid progradation clinoforms 200 2 m m 500 62 µ 200 2 m m 500 62 µ fig. 37. middle volgian palaeogeography (raukelv and lindemansbugt formations and correlatives) with sections through characteristic facies types and a block diagram showing the shelf-slope break in southern jameson land. note the different development of the southern platform-type basin and the northern half-graben basin. based on surlyk (1978b, 1990a) and surlyk et al. (1981); for legend, see fig. 6. beds. a system of large sandwaves driven by southwards flowing coast-parallel marine currents acted as a conveyor belt, which transported the coarse-grained pebbly sand to the front of the high-angle clinoforms forming the marine-reworked delta front (surlyk & noe-nygaard 1991). when the clinoform beds reached the shelf edge they prograded across the shelf slope break and down the upper slope where they started to collapse and the sand moved down the slope as sandy debris flows and high-density turbidity currents to be deposited on the slope and at the base-of-slope as thick massive sands forming the sjællandselv member of the hareelv formation (figs 4, 5, 40, 41). progradation continued, interrupted by slumping and development of debris flows, and on the upper slope all transitions between clinoform beds and massive slumped sands can be seen, whereas the base-of-slope is characterised by massive sandstones only. further out in the basin, the massive sands are interbedded with black mudstone; this facies belongs to the youngest part of the katedralen member of the hareelv formation. the clinoform beds are overlain by extensive sheets of glauconitic sandstone, up to 5 m thick, and in some cases by bioturbated highly fossiliferous mudstone. the sandstones were deposited slowly during transgression of the shelf-margin wedges. they are topped by sharp omission surfaces with great densities of diplocraterion habichi. the top of the raukelv formation (rauk plateau member) is incised by a network of highly sinuous, rather steep-walled sandstone-filled channels, which lead into a major valley at the delta front (figs 42, 43). the valley is about 10 km wide, up to 100 m deep and was probably enlarged by retrogressive slumping of the delta front. it is filled with an onlapping fully marine coarsening-upwards succession constituting the ryazanian hesteelv formation, consisting of black mudstones and heteroliths (crinoid bjerg member, p. maynci – h. kochi chronozones) passing up into extremely fossiliferous shelly sandstones (muslingeelv member, h. kochi –?s. analogus chronozones). other wide, deeply incised valleys occur at the same level in the southern outcrop area. 701 kap leslie fm parnasryggen mb hartz fjeld fm lower hennigryggen mb fig. 38. low-angle clinoform-bedded sandstone of the lower hennigryggen member (middle volgian), hartz fjeld formation, milne land. 702 e w s n shelf–slope break shelf slope fig.41 fig. 40. sandy shelf-margin wedge prograding across the shelf-slope break and down the slope to the left. note the overriding nature of succeeding shelf-margin sand bodies (top of section). regional dip 2–3°. inclination of slope to the left 10–20°. sandstone unit about 40 m thick. middle volgian fynselv member, raukelv formation, southern jameson land. fig. 39. shelf-margin wedge of the upper raukelv formation (rauk plateau member, middle volgian). it consists of coarse-grained pebbly sandstone showing high-angle clinoform bedding and occurs in a basinally-isolated position in south jameson land. 703 rauk plateau mb onlap channel incised into top surface of raukelv fm onlap sb top rauk plateau mb muslingeelv mb crinoid bjerg mb sb/base hesteelv fm n s fig. 41. slump scar overlain by massive sandstone. top of slope pictured on fig. 40. raukelv formation, southern jameson land. fig. 43. erosional unconformity at the middle volgian top of the raukelv formation formed by incision of a wide shelf-edge valley filled with marine siltstone (crinoid bjerg member, lower ryazanian) clearly wedging out towards valley margin (to the left) and overlain by highly fossiliferous sandstone (muslingeelv member, lower ryazanian). field of view c. 1 km. fig. 42. margin of highly sinuous tributary incised in middle volgian top of raukelv formation, draped with pebble veneer and filled with coarse-grained ryazanian sandstone. southern jameson land. sb, sequence boundary. the lower hennigryggen member (middle volgian) at the western basin margin in milne land is generally finer grained and has a higher diversity and density of marine trace fossils than the raukelv formation in the basin centre. the former is interpreted as highstand deposits, while the latter was deposited during a succession of major sea-level falls when coastal regressions reached the shelf-slope break close to the axis of the basin. there is thus a spatial separation between early highstand deposits and the associated late highstand–lowstand deposits in the order of 100 km. thin, widespread sheets of transgressive sandstones and mudstones drape the regressive shelf-edge sand bodies and probably extend further in a landwards direction. 704 bernbjerg fm payer dal fm pelion fm w fig. 44. westwards tilted jurassic fault block exposed on the south coast of kuhn ø. peneplaned basement directly overlain by shallow marine sandstones of the middle jurassic pelion formation succeeded by the backstepping middle–upper jurassic payer dal and bernbjerg formations (see figs 30, 31). the volcanic ridge is c. 1000 m high. basement fig. 45. coarse conglomerates and breccias of the middle volgian rigi member (lindemans bugt formation), wollaston forland, deposited on a fault scarp talus apron immediately adjacent to the dombjerg fault, western wollaston forland (fig. 2). view towards the north. tilted peneplain crystalline basement basalt e 705 correlative deposits are not known from the traill ø region but are well exposed in the wollaston forland area where they form the deep-water, conglomeratedominated wollaston forland group. this region underwent major rotational block faulting in the middle volgian resulting in the formation of deep-water halfgrabens (figs 3, 44). they were filled mainly by conglomerates, pebbly sandstones, sandstones and subordinate mudstones deposited by a variety of sediment gravity flows (figs 45, 46, 47; surlyk 1978b, 1984, 1989). the group is up to about 3 km thick in the deepest part of the hanging wall, adjacent to the main fault scarp, and decreases to 5–10 m over much of the eastern block crest. the dominant motif is a fining-upwards cycle, tens to some hundreds of metres thick, thought to reflect a major fault episode and associated deepening (fig. 47; surlyk 1978b). the cycles may correspond to the contemporaneous shallow marine sequences of the raukelv formation in jameson land. the group is subdivided into the volgian – upper ryazanian lindemans bugt formation, which is dominated by boulder conglomerates, and the finer-grained uppermost ryazanian – hauterivian palnatokes bjerg formation. deposition began with the formation of submarine talus breccias dominated by metre-sized blocks and rapid progradation of interbedded pebbly sandstones, sandstones and mudstones (laugeites ravine member). a temporal succession of environments developed ranging from submarine talus cones, through slope aprons to coalescent fan deltas – submarine fans (rigi member; surlyk 1989). the coarse-grained part of the system extends up to 25 km eastwards, away from the fault scarp, where it passes rapidly into black sandy mudstones (niesen member) close to the axis of the basin, at which point the gravity flows were checked by the opposing west-dipping hanging-wall slope. the group shows an overall fining-upwards trend, which culminated in an important late ryazanian drowning event. w e fig. 46. deep-water conglomerates deposited from sediment gravity flows in the rift-climax half-graben of the middle volgian lindemands bugt formation, wollaston forland. note clast imbrication indicating eastwards transport (arrow). 706 a b j2.6. late ryazanian – hauterivian: end of rifting and regional drowning hesteelv and upper hartzfjeld formations, upper niesen member (lindemans bugt formation) and palnatokes bjerg formation a regional flooding event took place in the latest ryazanian b. mesezhnikovi chron. in southern jameson land, the top of the raukelv formation was drowned, including the fills of the distributary channel network and the valleys incised in the delta front. massive gravity-flow sandstones directly overlie the pebbly clinoform bed of the rauk plateau member, which is the youngest shelf-margin wedge of the raukelv formation. the massive sandstones are interpreted as base-of-slope sandstones triggered from the front of latest ryazanian – early valanginian shelf-margin wedges, which occupied a more landwards position compared to the raukelv formation clinoform beds. the youngest shelf-margin wedges, however, are not preserved owing to modern erosion, and the massive base-of-slope sandstones form the youngest pre-quaternary deposits in jameson land. after a long period of basin margin emergence and erosion, the transgression reached milne land in the (?)early valanginian, and middle volgian highstand delta sandstones of the lower hennigryggen member are overlain by similar valanginian sandstones deposited during the regional drowning event. in the wollaston forland region, block-faulting and tilting waned, and the very coarse-grained deposits of the lindemans bugt formation are overlain by the finer grained palnatokes bjerg formation. deposition of proximal conglomerates, pebbly sandstones and sandstones (young sund member) persisted into the latest ryazanian – hauterivian, but mudstones dominate. oxygenation at the sea floor increased, and the coarse-grained deposits pass seawards into calcareous mudstones of the albrechts bugt member, which overlie the dark grey niesen member mudstones. the submerged crests of blocks became draped with claret-coloured, highly fossiliferous, bioturbated mudstones of the valanginian– hauterivian rødryggen member (figs 48, 49). the coarse-grained gravity flow deposits of the lindemans bugt formation were mainly transported downslope due eastwards, directly away from the fault scarps. with time, a more mature drainage and transport system developed. a transfer zone between two right-stepping border-fault segments between wollaston forland and th. thomsen land became an important entry point for coarse clastics, and the main transport became axial towards the south (young sund member). source areas and drainage patterns the main source area for the jurassic sedimentary basin complex was the greenland craton consisting of precambrian and caledonian metamorphic and intrusive rocks and thick proterozoic sedimentary successions of the hagen fjord and eleonore bay groups, which are exposed along the eastern margin of the craton. the crystalline basement of the liverpool land high constituted a minor source area east of the jameson land basin, but became gradually onlapped in the jurassic. continental devonian–carboniferous deposits were exposed north and north-west of jameson land in early jurassic time while areas north of jameson land were onlapped during the middle jurassic. the increased late jurassic rifting, block faulting and tilting resulted in uplift and erosion of footwall crests, which formed additional minor source areas. the main source terranes throughout the jurassic were thus weathered metamorphic basement and intrusive rocks, and, less importantly, devonian red conglomerates, sandstones and mudstones, carboniferous sandstones and mudstones. an additional source was formed by triassic conglomerates, sandstones and mudstones in the early jurassic due to uplift of the areas north of jameson land (surlyk 1977a, 1978a; surlyk et al. 1993; johnson & gallagher 2000). the drainage patterns underwent marked changes during the jurassic. the early jurassic was a time of tectonic quiescence and thermal subsidence following late permian – early triassic rift events. sedimentation was restricted to the jameson land platform, which was gently tilted towards the west. it was limited to the west by the n–s-trending stauning alper fault, which formed the border to the greenland craton. to the east, the platform was bordered by the peneplaned crystalline base707 facing page: fig. 47. large-scale fining-upwards succession in turbiditic deep-water conglomerates and sandstones of the middle volgian rigi member (lindemans bugt formation), wollaston forland. palaeotransport was towards the east, into the photographs. a: lower conglomerate-dominated part of the section. b: succeeding sandstone-dominated part. the encircled sandstone block is common to both photographs. 708 ment of the liverpool land high, which was limited to the east by an important fault, the liverpool land escarpment of larsen (1984), and to the west by a major fault running close to the east coast of hurry inlet, through klitdalen and further north through carlsberg fjord, where it passed into nneand nnw-trending splays. the land areas bordering the jameson land platform were relatively low-lying with a gentle relief. the drainage was centripetal and material was transported to the basin from the west, north and east (dam & surlyk 1998). onset of rifting in the late bajocian resulted in a marked change in drainage pattern. rifting was assoe basalt sbsb sb lower cretaceous bernbjerg fm jakobsstigen mb jakobsstigen fm pelion fm pelion fm barremian albrechts bugt mb albrechts bugt mb rødryggen mb fault stratigraphic boundary sequence boundarysb fig. 48. tilted jurassic fault block, south side of the valley of cardiocerasdal, southern wollaston forland. the shallow marine middle jurassic pelion formation is exposed at the base of the succession and is overlain by a backstepping succession of the upper jurassic jakobsstigen and bernbjerg formations. this early-rift succession is faulted down to the east; to the west, it is unconformably overlain by valanginian–(?)hauterivian late-rift conglomerates (cgl) and sandstones (sst) of the young sund member, whereas to the east on the more crestal area it is overlain by the roughly contemporaneous light grey mudstones of the albrechts bugt member and red mudstones of the rødryggen member. the whole syn-rift succession is truncated by an irregular regional unconformity, and overlain by lower cretaceous (barremian) post-rift mudstones. mid-cretaceous sbrødryggen mb sb angularunconformity bernbjerg fm e w fig. 49. black kimmeridgian bernbjerg mudstones unconformably overlain by red valanginian–hauterivian deep-water mudstones of the rødryggen member, which is in turn unconformably overlain by dark-grey barremian post-rift mudstones. locality to the left in fig. 48. w jakobsstigen fm jakobsstigen fm pelion fm pelion fm pelion fm barremian sst sst albrechts bugt mb albrechts bugt mb young sund mbyoung sund mb bernbjerg fm jakobsstigen fm cgl cgl 709 ciated with westwards tilting of fault blocks bordered by mainly n–s-trending faults. the basin was extended far north of the jameson land platform and two elongate embayments were formed with their heads in the clavering ø and hochstetter forland areas. the drainage was longitudinal with the main entry points at the heads of the embayments, which can be considered tectonic estuaries. the large trunk rivers appear to have followed the tectonic grain and a major river system thus probably flowed northwards or southwards west of the depositional basin along the axes of the blocks bordered to the east by the stauning alper fault – post-devonian main fault (compare figs 2, 16 and 17). the fluvial systems could only enter the depositional basins in transfer or relay zones where the border faults side-stepped en echelon. sediment transport within the basin was now almost exclusively axial towards the south as reflected by palaeocurrent data, southwards grain-size decrease and facies changes. the northwards extension of the depositional basin resulted in draping of most of the late palaeozoic source lands and this effect was increased by late middle – late jurassic eustatic sea-level rise and associated flooding of former land areas. only the northwestern part of the devonian basin was probably still exposed and may have acted as a source area in addition to the crystalline basement rocks of the greenland craton and the liverpool land high. the increase in late jurassic rifting activity, in concert with sea-level rise, led to further compartmentalisation of the rift complex, and deposition became dominated by black muds and local massive sands; most of the coarser fluvial sediment load was probably trapped in estuaries. impressive fault scarps were formed during the volgian rift climax, especially in the wollaston forland region (vischer 1943; maync 1947, 1949; surlyk 1978b, 1984, 1989). the scarps were sources for coarse breccias and conglomerates, deposited as transverse fault-scarp aprons and coalescent fans and axial basin floor fans (surlyk 1978b, 1984). as rifting waned, the drainage areas enlarged and older, probably carboniferous, conglomerates were reworked and shed into the half-graben at transfer or relay zones formed between segments of the main fault system. the east greenland rift complex thus witnessed longterm change in borderland topography and drainage pattern. early jurassic centripetal drainage of relatively low-lying borderland by numerous smaller rivers changed in middle and late jurassic times into a system of major n–s-running trunk rivers with deltas at transfer or relay zones between the main n–s-trending fault zones. the provenance areas at the same time 710 viking graben east greenland relative position of the shoreline basinwards landwards se ri es ju ra ss ic st ag es vo lg ia n k im m er id gi an o xf or di an c al lo vi an ba th on ia n ba jo ci an a al en ia n to ar ci an chronozones s. preplicimphalus s. lamplughi s. primitivus t. oppressus t. anguiformis g. kerberus g. okusensis g. glaucolithus p. albani v. fittoni p. rotunda p. pallasioides p. pectinatus p. hudlestoni p. wheatleyensis p. scitulus p. elegans a. autissiodorensic a. eudoxus a. mutabilis r. cymodoce p. baylei a. rosenkrantzi a. regulare a. serratum a. glosense c. tenuiserratum c. densiplicatum c. cordatum q. mariae q. lamberti p. athleta e. coronatum k. jason s. calloviense m. macrocephalus c. discus o. aspidoides p. hudsoni m. morrisi t. subcontractus p. progracilis a. enuiplicatus z. zigzag p. parkinsoni s. garanitiana s. subfurcatum s. humphriesianum e. sauzei w. laeviuscula h. discites g. concavum l. murchisonae l. opalinum d. levesquei g. thouarsense h. variabilis h. bifrons h. falciferum u pp er m id dl e lo w er u m l u m l u m l u l fig. 50. relative sea-level curves or, more precisely, curves showing landwards/basinwards shifts in the position of the coastline. north sea sea-level curve and stratigraphy after sneider et al. (1995). note the remarkable similarity between this curve and the east greenland curve based on the stratigraphy of jameson land and milne land. changed from basement and sediment to mainly basement-dominated. comparison with the north sea and adjacent areas jurassic basin evolution is remarkably similar in the north sea – north atlantic region in an area stretching from northern east greenland in the north to the central graben and the danish basin in the south (figs 1, 4, 50, 51). major changes in facies and depositional environments occur towards the south, however, in onshore england and northern france although the dutch north sea sector still shows marked similarities with the regions further north (see herngreen et al. 2003, this volume). the uniform development of the jurassic succession in east greenland and the west norway shelf is not surprising as the two areas essentially formed the western and eastern halves of the same basin complex until the onset of sea-floor spreading at the paleocene–eocene boundary (see doré 1992 for an overview). thus stewart et al. (1995, p. 321) in their description of the horda platform, offshore south-west norway, wrote “it is noteworthy, however, that the sequence architecture of similar middle to upper jurassic shallow marine deposits in east greenland is almost identical (surlyk 1991[a]), suggesting the operation of large-scale, regional processes”. a detailed comparison of the upper jurassic in east greenland (based on surlyk 1991a) and the north sea is strongly suggestive of the isochronous nature of maximum flooding surfaces (partington et al. 1993a, b). most of the sequence boundaries in the kimmeridgian – lower volgian of east greenland seem to correlate well with those identified within the kimmeridge clay formation in the wessex–weald basin by taylor et al. (2001). in some cases the biostratigraphic correlation is, however, not sufficiently precise mainly due to faunal provincialism. the overall evolution in the north sea – north atlantic region is similar in terms of tectonism and stratigraphy; there are, however, numerous smaller scale differences in timing of events and facies architecture. this may reflect real differences for example in timing of the onset and duration of major rift events and their interplay with eustatic changes in sea level, as well as the nature of source areas, drainage patterns and sediment influx. it may, however, also reflect lack of precision in biostratigraphic age determination, faunal provincialism in middle and latest jurassic times and use of different zonal schemes based on ammonites and dinoflagellate cysts, respectively. the north sea dinocyst zonation is also less wellcorrelated with standard ammonite zones than is commonly implied by the use of ammonite zonal names for flooding surfaces that are dated exclusively by dinocysts. it is thus noteworthy that the greatest similarity with the mainly ammonite-dated sedimentological and stratigraphic evolution of east greenland is shown by the onshore succession exposed along the moray firth, scotland which is also constrained by a detailed ammonite geochronology. precise correlation with the norwegian shelf successions and the north sea is more difficult. mismatches in the age of major sequence stratigraphic key surfaces, onset of tectonism and duration of stratigraphic events may or may not be real. the great similarity between the jurassic of east greenland, the north sea and norwegian shelf is welldemonstrated by the work of rattey & hayward (1993) and sneider et al. (1995) on the middle–upper jurassic of the north sea and viking graben, respectively, and by stewart et al. (1995) on the horda platform (fig. 4). jameson land is essentially a platform-type basin, which shows great similarity with the horda platform while the wollaston forland is a classical area for study of deepwater half-graben development and is an excellent analogue for the viking graben and the moray firth. relative sea-level curves for the viking graben and jameson land – milne land are remarkably similar, even allowing for uncertainties in dating, correlation and interpretation (fig. 50). the lower jurassic is not preserved or is poorly known over much of the north sea area due to erosion in connection with the development of the baselate aalenian unconformity. the lower jurassic of jameson land, east greenland, is however an excellent analogue for the lower jurassic of the conjugate norwegian margin. the remarkable similarity of the successions of the two areas allowed dam & surlyk (1995, 1998) to undertake a detailed sequence stratigraphic correlation. this has been corroborated and refined recently by a sequence stratigraphic interpretation of the uppermost ostreaelv formation – sortehat formation (hansen 1999). the jurassic of the inner moray firth is selected for a more detailed comparison with east greenland because both areas include outcrops and are dated mainly by ammonites (fig. 51). the jurassic onshore and offshore moray firth has recently been interpreted within a sequence stratigraphic framework (stephen et al. 1993; stephen & davies 1998) while much detail on the actual sections is compiled in an excellent field guide (trewin & hurst 1993). 711 comparison with the middle–upper jurassic of the moray firth the triassic–jurassic transition was marked by alluvial deposition both in east greenland and the moray firth, where the late rhaetian – sinemurian successions are represented by the kap stewart group and the dunrobin bay formation, respectively. the tops of both units are major sequence boundaries. marine transgression took place close to the sinemurian–pliensbachian boundary with maximum flooding in the lowermost pliensbachian p. taylori subchronozone of the u. jamesoni chronozone in both areas (sellwood 1972; dam & surlyk 1998). interestingly, a similar succession of events has recently been demonstrated from the island of bornholm in the baltic (surlyk et al. 1995; donovan & surlyk 2003, this volume). the remaining lower jurassic stages are not represented in the moray firth due to non-deposition and erosion associated with uplift of the north sea dome and formation of the mid-aalenian unconformity (the ‘mid-cimmerian unconformity’ of underhill & partington 1994, and many others). the erosional topography was gradually onlapped in late middle jurassic time by the brora coal formation. a similar development can be seen in milne land and in the areas further north in east greenland (figs 4, 25, 31) where middle and upper jurassic sandstones of the charcot bugt, pelion and payer dal formations onlap triassic, upper permian or crystalline basement (surlyk 1977a, 1978a). there is thus abundant evidence for major pre-middle jurassic uplift or doming in large parts of the north sea – north atlantic region. the age of the terrestrial and lagoonal brora coal formation is poorly constrained but it is roughly contemporaneous with the pelion formation and likewise represents the lower part of a long-term middle jurassic regressive– transgressive cycle (fig. 51). the top of the brora coal formation was deposited in a lagoonal environment followed by a short regressive interlude at the bathonian– callovian boundary represented by the brora coal. this regressive event is contemporaneous with a similar event in east greenland represented by the parnas member at the top of the pelion formation, which interrupts the long-term backstepping of the pelion–fossilbjerget couplet (fig. 4). the brora coal formation is overlain by the transgressive brora roof bed succeeded by mudstones and shales of the brora shale and brora brick clay with a thin intercalation of the very fine-grained glauconitic sandstone. this interval correlates with the goniomyakløft member (fossilbjerget formation), which represents the final marine drowning of the shallow marine sandstones of the pelion formation. maximum flooding took place in e. coronatum – earliest p. athleta chron time, at the middle–late callovian boundary, in both areas. maximum flooding was followed by progradation, in the moray firth represented by the upper callovian fascally siltstone – fascally sandstone – clynelish quarry sandstone package which corresponds to the equally progradational athene member of the olympen formation (fig. 51). the latter unit is dominated by massive sandstones deposited from sediment gravity flows on the slope of a high-angle clinoform bedded shelfmargin wedge. the massive, wavy-bedded clynelish sandstone is clearly of similar nature, and was deposited from sediment gravity flows. this interpretation contrasts with previous interpretations, which involve tidal influence (hurst 1993; stephen et al. 1993; stephen & davies 1998). the alternative interpretation of depositional processes and environment offered here casts some doubts on the reality of the sequence stratigraphic interpretation of stephen et al. (1993) and stephen & davies (1998). the sharp base of massive gravity flow sandstones may reflect the nature of the sedimentary processes rather than representing a regressive surface of erosion caused by sea-level fall. the base of the clynelish quarry sandstone may thus be strongly diachronous reflecting repeated collapse of sand along the shelf edge and downslope transport of the triggered gravity flows (unpublished data, f. surlyk 2002). the clynelish quarry sandstone is overlain by a thin fine-grained unit, which is a correlative of the shales of the hades member (olympen formation); it represents the drowning of the first shelf-edge delta and associated slope deposits at the callovian–oxfordian boundary. in the moray firth, the drowning interval was succeeded by renewed progradation of the brora sandstone member. the member is of early–middle oxfordian age but the dating is not well-constrained. the much-figured but somewhat enigmatic high-angle clinoform-bedded sandstones at strathsteven cliffs are referred to the brora sandstone member and have mainly been interpreted as having been deposited by migrating sand dunes in a strongly tidally influenced sea (e.g. stephen & davies 1998, fig. 15) or a large seawards-dipping bar structure (sykes 1975; hurst 1993). the clinoform beds are here interpreted to represent a shelf-margin wedge or delta due to the remarkable similarity to the shelf-margin wedge of the contemporaneous zeus member, olympen formation, and even more to the stacked volgian shelf-margin wedges of the raukelv formation which also migrated seawards at a 712 right angle to the basin-bounding fault during periods of sea-level fall. the upper callovian – middle oxfordian of both areas thus shows a highly similar development both stratigraphically and in terms of depositional processes and environments. as interpreted here, both areas show a development from a ramp to a shelf-break basin associated with forestepping late callovian and early–middle oxfordian sandy shelf-margin wedges separated by a latest callovian – earliest oxfordian drowning event. when the deltas reached the shelf-slope break in both areas, the delta front collapsed and sands were transported down the slope by sediment gravity flows (surlyk & noe-nygaard 2001b; larsen & surlyk 2003, this volume). the middle oxfordian drowning of the brora sandstone was accompanied by a marked change in facies 713 zeus mb hades mb athene mb goniomyakløft mb parnas mb olympen fm fossilbjerget fm backstepping pelion fm fossilbjerget fm sb ds ds ds dsds ds ds ds sb sb mfs o xf or di an c al lo vi an ba th on ia n upper upper lower middle lower– middle br or a c oa l br or a a rg ill ac eo us br or a a re na ce ou s doll inverbrora brora shale fascally siltstone brora sandstone s. calloviense p. athleta q. lamberti mid stage chronozone formation member c. densi. bal. clyn. quarry sst arda. lst fa. sst brora brick clay glauc. sst e. coronatum k. jason m. macro. c. discus brora roof bed brora coal 50 inner moray firth jameson land 0 m 100 150 200 coal claystone mudstone heterolithic (sst/mst) sandstone carbonate bands/concretions fig. 51. comparison of the middle – lower upper jurassic stratigraphies of the inner moray firth (modified from trewin & hurst 1993; stephen & davies 1998) and jameson land. note the remarkable similarity in overall trends and age of main flooding events. the jameson land schematic section is not to scale, thicknesses being normalised to the moray firth section. c. densi., c. densiplicatum; m. macro., m. macrocephalus; arda., ardassie; bal., balintore; clyn., clynelish; fa., fascally; glauc., glauconitic; ds, drowning surface; mfs, maximum flooding surface; sb, sequence boundary. and rapid deepening of the basin. upper oxfordian deposits are only known from offshore moray firth where they are represented by a backstepping progressively finer grained succession with peak transgression in the latest kimmeridgian or earliest volgian (stephen & davies 1998). this is similar to the development in east greenland where drowning of the shallow marine olympen, charcot bugt and jakobsstigen formations was succeeded by a backstepping succession of increasingly finer grained units (figs 4, 30, 31). the upper oxfordian – kimmeridgian interval records the onset of the main rifting phase with tilting of fault blocks and extensional half-graben development in the moray firth area (underhill 1991). the coastal outcrops along the helmsdale fault present a variety of sediment gravity flow deposits ranging from sandstone breccias and quartz sandstones of the allt na cuile formation to conglomerates and breccias of the kintradwell and helmsdale boulder beds (pickering 1984; macdonald & trewin 1993; wignall & pickering 1993; theriault & steel 1995). this succession is again remarkably similar in tectonic setting and facies to the thick deep-water conglomerate-dominated wollaston forland group of northern east greenland (surlyk 1978b, 1984, 1989), as originally noted by b.w. sellwood (in: anderton et al. 1979). accumulation of fault-scarp conglomerate aprons in the moray firth started, however, in the earliest kimmeridgian and culminated with deposition of the helmsdale boulder beds in the late kimmeridgian – middle volgian. the wollaston forland group is of early or middle volgian to hauterivian age with a climax of coarse-grained deposition in the middle volgian. the oldest rift climax deposits, the allt na cuile formation, occur immediately adjacent to the scarp of the helmsdale fault and are represented by immature turbidites and related deposits (wignall & pickering 1993; damholt 1996). in the most distal exposures, the allt na cuile formation comprises massive sandstones, which have undergone post-burial liquefaction and intrusion into the adjacent sediments (unpublished data, f. surlyk 2002) and this part of the formation shows great similarity to the extremely well-exposed hareelv formation in jameson land, east greenland (surlyk 1987; surlyk & noe-nygaard 1995, 2000b, 2001b). the jurassic sections of the two areas thus include an upper rhaetian – lower jurassic pre-rift succession, which can be considered a late post-rift succession with respect to a major late permian – early triassic rift event. it shows an overall long-term transgressive development, which was terminated by widespread doming or regional uplift in late early jurassic time. subsidence and onlap of the uplifted areas started in the late aalenian in the moray firth – north sea region and in the late bajocian in east greenland. the subsequent development records incipient rifting in the bajocian– bathonian, followed by increasing rifting and progressive drowning of the sand-dominated systems in late bathonian – callovian times. maximum drowning took place close to the middle–late callovian boundary with a subsequent change from a ramp to a shelf-break basin. sandy shelf-margin wedges were formed in the late callovian and early–middle oxfordian, separated by a drowning event at the callovian–oxfordian boundary. massive sands were shed from the delta fronts and deposited on the slope by sediment gravity flows. a major phase of basin reorganisation and facies change took place in the late oxfordian when the sandy shelfmargin wedges were drowned. this event records the onset of major rifting, which continued through the kimmeridgian to culminate in the middle volgian. this general theme can be recognised throughout the north sea – north atlantic region. the similar development of the jurassic in the north sea – north atlantic region has two basic implications. one is that outcrop analogues exist in east greenland for practically all types of offshore petroleum reservoirs and play types, allowing study of dimensions, 3-d geometry, facies changes and predictability of major reservoir types. the other is that not only the similarities but also the differences in development allow conclusions to be drawn concerning factors governing basin evolution and sedimentation. summary and conclusions the rhaetian–hauterivian succession of east greenland comprises two megasequences, j1 and j2, representing different tectonic regimes. the rhaetian – lower bajocian j1 package forms a pre-rift megasequence with respect to the overlying syn-rift deposits. genetically it is more precisely described as a post-rift package with respect to late permian – earliest triassic and late early triassic rift events. the upper bajocian – hauterivian package forms a syn-rift megasequence, formed during a protracted rift episode that began in the late bajocian, increased through the bathonian–kimmeridgian, culminated in the volgian and waned in the latest ryazanian – hauterivian. eight low-order tectonostratigraphic sequences are recognised in the rhaetian–hauterivian succession, two in the pre-rift megasequence (j1.1–1.2) and six in the 714 syn-rift megasequence (j2.1–2.6). they correspond to major changes in tectonic style, basin configuration, drainage pattern, transport directions and depositional environments. their boundaries are major unconformities or drowning surfaces, and they are not depositional sequences in the sense of posamentier & vail (1988) but rather tectonostratigraphic sequences in the sense of surlyk (1991b). sequence stratigraphic interpretation of the succession illustrates the difficulties in using one, simple, standardised sequence stratigraphic concept or method. unconformities are well-developed in some parts of the succession and very difficult or impossible to recognise in others. this is also the case with drowning surfaces and marine transgressive or regressive erosion surfaces. the concept of a maximum flooding surface is sedimentologically meaningless but intervals corresponding to maximum flooding can be identified in the lower part of many mudstone-dominated coarsening-upwards units. they are not considered of major importance in most high-resolution outcrop studies but have their main importance in subsurface studies, commonly of lower resolution. parasequence stacking patterns are well-developed at some levels and completely missing at others. the rhaetian–sinemurian part of the pre-rift megasequence (j1.1) is fluvio-lacustrine and a large number of high-order sequences are recognised, interpreted as having been caused by climatically-controlled lake-level changes, possibly within the milankovitch frequency band. long-term lake-level changes may, on the other hand, reflect base-level changes controlled by fluctuations in sea level. the pliensbachian – lower bajocian shallow marine – paralic part of the pre-rift megasequence (j1.2) shows unconformities and ravinement surfaces, whereas welldeveloped stacking patterns are rarely developed. sequences are accordingly mainly well-defined, whereas systems tracts are difficult to identify. the upper bajocian – lower kimmeridgian part of the syn-rift megasequence (j2.1–2.4) was deposited during a period of major expansion and reorganisation of the basin. deposition was fully marine with the exception of basal fluvial successions that are developed in a few areas. stacking patterns are well-developed, but differentiation between unconformities, ravinement and drowning surfaces is difficult. it is thus in many cases an open question as to whether the basic coarseningupwards units are parasequences or simple sequences, or in other words if the top surfaces reflect drowning alone or sea-level fall followed by flooding. rapid axial southwards progradation of an extensive shallow marine sandy shelf and shoreface system took place in late bajocian time. deposition was aggradational until the middle–late bathonian when largescale backstepping started, culminating in complete drowning of the sandy system in the middle callovian punctuated by a minor progradational episode at the end of the early callovian. backstepping was probably due to increased rates of rifting overprinted by middle–late jurassic eustatic sea-level rise. the jameson land basin changed from a ramp to a shelf-break margin in the late callovian, probably caused by intensified rifting. two regional progradational events took place in the latest callovian – middle oxfordian marked by the incoming of base-of-slope sands and sandy shelf-margin wedges in the jameson land area, by deposition of shallow marine sands on traill ø, and alternating shallow marine sands and thin coastal plain muds in wollaston forland. drowning of the middle oxfordian shallow marine deposits was followed by major regional transgression and deepening, punctuated by a short-lived late oxfordian progradational event in milne land. by late kimmeridgian time, the east greenland basin was characterised by deposition of mud in fairly deep water under poorly oxygenated conditions. the deepest part of the basin in jameson land received sandy gravity flow deposits throughout the late oxfordian – volgian. the sands underwent post-burial liquefaction and were injected into the surrounding mudstones, forming an impressive intrusive sand-body complex. sand remobilisation and intrusion were probably triggered by riftassociated earthquakes. the volgian rift-climax succession is differently developed in the southern and northern parts of the basin complex, reflecting different degrees of block tilting. the southern part centred over jameson land behaved as a wide, coherent platform and underwent relatively minor tilting. maximum early kimmeridgian drowning was thus followed by rapid volgian progradation and regression. sand deposition took place at the western basin margin during highstands and shifted to the shelf edge located close to the basin axis during late falls and lowstands, where thick clinoform beds were formed. reworked transgressive drapes were deposited during succeeding sea-level rises. highstand and lowstand deposits of the same sequence are thus separated by many tens of kilometres. in the northern part of the basin complex, in the wollaston forland area, deep half-grabens were formed during the middle volgian rift climax, and thick con715 glomerates, pebbly sandstones, sandstones and minor mudstones were deposited off the fault scarp from sediment gravity flows. the main motif is a fining-upwards unit, tens of metres thick, thought to reflect major phases of down-faulting and deepening. sequence stratigraphic key surfaces cannot be identified due to the deep-water nature of the succession but the fining-upwards units may be correlatives of the contemporaneous shallow marine sequences in the jameson land area. rifting waned in ryazanian–hauterivian times and a regional drowning event took place in the latest ryazanian. in jameson land, this is marked by the incoming of massive base-of-slope sands overlying sandy, shallow marine clinoform-bedded shelf-margin wedges. in the wollaston forland area, deposition of deep-water conglomerates and pebbly sandstones was succeeded by much finer grained sediments, although coarse clastic deposition still took place. the drainage pattern became more mature and the main sediment entry point was at a transfer zone between right-stepping en echelon border faults. the sediment gravity flows changed from a mainly lateral eastwards to an axial southwards transport direction. the more distal deposits are dominated by calcareous mudstones, and red mudstones drape the submerged block crests. the rhaetian–hauterivian succession of east greenland thus provides a well-exposed record of rhaetian – early bajocian pre-rift deposition (j1). this was succeeded by late bajocian – hauterivian syn-rift deposition (j2) characterised by different degrees of block faulting and tilting in different parts of the basin. the jurassic of east greenland shows marked similarities to the west norway shelf, the north sea and in particular to the inner moray firth. differences in timing of events may be real or reflect inaccuracies in biostratigraphic correlation due to faunal and floral provincialism, poor resolution of dinocyst stratigraphy in several intervals and poor correlation between ammonite and dinocyst zonations. acknowledgements i am grateful to the danish natural science research council, the carlsberg foundation, norsk hydro, british petroleum, conoco, statoil, saga, norwegian petroleum directorate, amoco and the danish ministry of the environment and energy for economic support over the years for field work and ph.d. stipends to a number of students. i thank peter alsen for new information on the age of the rødryggen member, john h. callomon and gregers dam for constructive reading of the manuscript, tony hallam and bruce w. sellwood for useful reviews, louise a. hansen and christian hagen for drafting, merete vesterager for word processing, ole b. berthelsen for dark-room work and jon r. ineson for careful editing. references alsen, p. & surlyk, f. in press: maximum middle jurassic transgression in 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(eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 75–114 (this volume). ziegler, p. 1988: evolution of the arctic – north atlantic and the western tethys. american association of petroleum geologists memoir 43, 198 pp. manuscript received 19 june 2001; revision accepted 27 september 2001. chronostratigraphyma stage chronostratigraphy hauterivian valanginian ryazanian volgian berriasian tithonian kimmeridgian oxfordian callovian bathonian bajocian aalenian toarcian pliensbachian sinemurian hettangian rhaetian major extension on atlantic margin magnus/møre relative tectonic quiescence in northern north sea onset of major extension in central north sea onset of major extension in northern north sea major extension in nns regional thermal uplift thermal subsidence after late permian – triassic rifting n.sea sequence tectonic controls on relative sea level northern north sea (nns) central north sea (cns) north viking graben horda platform jameson land milne land wollaston forland – kuhn ø east greenl. tect. strat. seq. c re ta ce ou s ju ra ss ic tr ia ss ic lo w er lo w er u pp er m id dl e u pp er k20 k10 j70 j60 j50 j40 j30 j20 j10 j0 highly condensed or absent due to submarine non-deposition basin basin shelf shelf shelf paralic draupne fm heather fm sognefjord fm ? fensfjord fm krossfjord fm tarbert fm rannoch/etive fms ness fm drake fm cook fm amundsen fm statfjord fm ra he z at go focondensed condensed so sk as e ræ alb rh ho gu pr i i lake lagoon lake delta pr delta har ha s pa pe u k h g a m c ko h r falpal y li ri n b j pe ug pay mu ba onlaps crystalline basement onlaps crystalline basement onlaps upper permian s n w e w e s n 2.6 2.5 2.4 2.3 2.2 2.1 1.2 1.1 u m l u l u m l u m l u m l u l u m l u l u l u l hauterivian valanginian ryazanian volgian berriasian tithonian kimmeridgian oxfordian callovian bathonian bajocian aalenian toarcian pliensbachian sinemurian hettangian rhaetian c re ta ce ou s ju ra ss ic tr ia ss ic lo w er lo w er u pp er m id dl e u pp er u m l u l u m l u m l u m l u l u m l u l u l u l 210 200 190 180 170 160 150 140 130 ? ? ? a al alb as at b ba c e f fo g go gu aldinger elv mb albrechts bugt mb albuen mb astartekløft mb athene mb bernbjerg mb bastians dal fm charcot bugt fm elis bjerg mb falskebugt mb fossilbjerget fm gråkløft mb goniomyakløft mb gule horn fm h ha har he ho i j k ko l li lnh m hartz fjeld fm hades mb hareelv fm hesteelv fm horsedal mb innakajik fm jakobsstigen fm krebsedal mb kosmocerasdal mb laugeites ravine mb lindemans bugt fm lepidopteriselv, nathorst fjeld, harris fjeld mbs mudderbugt mb mu n o pa pal pay pe pr r ra rh ri ræ s muslingebjerg fm niesen mb ostreaelv fm parnas mb palnatokes bjerg fm payer dal fm pelion fm primulaelv fm rødryggen mb raukelv fm rhætelv fm rigi mb rævekløft fm sjællandselv mb sk so t u ug y z skævdal mb sortehat fm trefjord bjerg mb ugleelv mb ugpik ravine mb young sund mb zeus mb coal mass flow sandstone source rock lacustrine deltas, sand-dominated open lake mudstones alluvial/delta plain – paralic, sand-dominated lagoonal sandstones/mudstones shallow marine sandstones shelf transition-zone siltstones/heteroliths marine shelf/basinal mudstones deep marine sandstones deep marine conglomerates marine argillaceous limestones red marine mudstones hiatus/condensed lnh t o east greenland lithostratigraphic units l fig. 4. jurassic stratigraphic development in the north sea in general (modified from rattey & hayward 1993; rhaetian–hettangian based on partington et al. 1993a, b), north viking graben (modified from sneider et al. 1995), horda platform (modified from stewart et al. 1995) and jameson land, milne land and wollaston forland in east greenland. note the similarity in middle jurassic onlap and progressive middle–late jurassic backstepping, but note also the differences in timing of events, which may be real and/or due to poor dating and difficulties in north–south correlation and between ammonite and dinocyst zonations. e2019430202-01 the greenland ice sheet has been losing mass in response to increased surface melting (khan et al. 2015; van den broeke et al. 2017) as well as discharge of ice from marine terminating outlet glaciers (van den broeke et al. 2009; box et al. 2018). marine terminating outlet glaciers flow to the ocean where they lose mass by e.g. iceberg calving. currently, the mass loss from the greenland ice sheet is the largest arctic contributor to global sea-level rise (van den broeke et al. 2009, 2017; box et al. 2018). therefore, monitoring changes in the greenland ice sheet is essential to provide policy makers with reliable data. there is a consensus that most marine terminating outlet glaciers have retreated in recent decades, and that the increased calving rates are a response to recent atmospheric and oceanic warming (e.g. box et al. 2018; moon et al. 2018). the rate of dynamic mass loss is determined by changes of the glacier calving front (i.e. its terminus) position, ice thickness and changes in ice flow. ocean temperature and fjord circulation also influence the calving front stability by melting the glacier below the water line, thinning the ice that is in contact with water (moon et al. 2014). change in calving front position is therefore an important indicator for monitoring the dynamic behaviour of the upstream area of the ice sheet, which is further modulated by local topographic features and buttressing effects (rignot & kanagaratnam 2006; nick et al. 2009). the programme for monitoring of the greenland ice sheet (promice) is dedicated to monitoring changes in the mass budget of the greenland ice sheet, including monitoring of the calving front lines of marine terminating outlet glaciers. here, we present an updated collection of annual measurements of end-of-melt-season calving front lines for 47 marine terminating outlet glaciers in greenland between 1999 and 2018. we also present an example application of update of annual calving front lines for 47 marine terminating outlet glaciers in greenland (1999–2018) jonas k. andersen1, robert s. fausto*1, karina hansen1, jason e. box1 and the promice project team** **signe b. andersen1, andreas p. ahlstrøm1, dirk van as1, michele citterio1, william colgan1, nanna b. karlsson1, kristian k. kjeldsen1, niels j. korsgaard1, signe h. larsen1, kenneth d. mankoff1, allan ø. pedersen1, christopher l. shields1, anne solgaard1 and baptiste vandecrux1 geus bulletin is free to individuals and institutions in electronic form. the author(s) retain copyright over the article contents. research article | open access geus bulletin vol 43 | e2019430202 | published online: 26 june 2019 https://doi.org/10.34194/geusb-201943-02-02 daugaard jensen kangerdlugssuaq fenris helheim midgaard tingmjarmiut ikertivaq qajuuttap sermilik kangiata nunata sermia jakobshavn sermeq avannarleq store lille kangigdleqsermeq silardleq perdlerfiup sermia kangerdluarssup sermia rinkumiamako ingia upernavik nunatakavsaup sermia nunatakassaap sermia hayes steenstrup kong oscars docker smith humboldt petermann nioghalvfjerdsfjorden zachariae storstrømmen academy ryder hagen ostenfeldsteensby fig. 1. location of each of the surveyed outlet glaciers (table 2). https://doi.org/10.34194/geusb-201943-02-02 e2019430202-02 the data set, in which we estimate area changes for this group of glaciers since 1999. the greenland calving front lines were measured from optical satellite imagery obtained from landsat, aster, and sentinel-2 (table 1). the promice calving front product is freely available for download as esri shapefiles (https://doi.10.22008/promice/data/calving_front_lines). methodology calving front lines were digitised through manual delineation in optical satellite images at the end of melt season every year from 1999 to 2018. the end of melt season was determined for each glacier by comparing a series of images from july–november (dependent on latitude) and selecting the one in which the glacier is at its minimum position. prior to 2018, calving front lines were digitised primarily using landsat-7 and landsat-8 images. in 2018, front lines were digitised solely using sentinel-2 imagery. we used mainly the panchromatic band to identify front line positions in landsat-7 images, a combination of three visible/nir bands in landsat-8 images and the nir band in sentinel-2 images (table 1). the resulting product is a collection of shapefiles, containing a polyline for each glacier for each year in the measurement period (1999–2018). the product comprises 47 of the largest greenland marine terminating outlet glaciers in the standard wgs-84 (epsg:4326) projection (see the glacier location map; fig. 1 and table 2). while the total number of glaciers is arbitrary, the ensemble of glaciers was selected to include the largest marine terminating outlet glaciers (measured in width) and to comprise glaciers from all along the greenland coast. table 1: overview of images used in the mapping of calving front lines satellite sensor band resolution image coverage (%) landsat 5 tm 3 30 m 0.8 landsat 7 etm+ 8 15 m 62.1 landsat 7 etm+ 3–5 30 m 4.9 landsat 8 oli 8 15 m 0.4 landsat 8 oli 4–6 30 m 22.0 terra aster 1–3 15 m 4.5 sentinel-2 msi 8 10 m 5.3 image coverage indicates the number of images used from each sensor as a percentage of the total number of images used. auxiliary images used in cases of distortion due to clouds or missing data bands in landsat 7 images are not counted. landsat images can be obtained from the usgs (https://earthexplorer.usgs.gov/), while sentinel-2 data can be obtained from the copernicus open access hub (https://scihub.copernicus.eu/). image coverage indicates the number of images used from each sensor as a percentage of the total number of images used. auxiliary images used in cases of distortion due to clouds or missing data bands in landsat 7 images are not counted. landsat images can be obtained from the usgs (https:// earthexplorer.usgs.gov/). sentinel-2 data can be obtained from the copernicus open access hub (https://scihub.copernicus.eu/). glacier name lat. long. width net area average area (°n) (°e) (km) change change (km2) (km2 a-1) zachariae 78.90 –20.14 24.6 –409.6 –21.6 petermann 80.10 –61.17 17.4 –289.5 –15.2 humboldt 79.50 –64.61 89.0 –259.4 –13.7 hagen 81.53 –28.50 9.1 –172.8 –9.1 jakobshavn 69.18 –49.73 11.4 –137.0 –7.2 storstrømmen 76.71 –22.47 31.9 –99.5 –5.2 nunatakassaap sermia 74.62 –56.34 5.4 –70.5 –3.7 steensby 81.20 –53.90 4.5 –69.5 –3.7 ostenfeld 81.60 –45.20 7.0 –65.0 –3.4 79 fjorden 79.60 –20.17 42.2 –51.7 –2.7 steenstrup 75.28 –57.89 16.2 –50.8 –2.7 kangerdlugssuaq 68.61 –32.93 6.0 –45.9 –2.4 midgaard 66.45 –36.73 3.8 –40.9 –2.2 upernavik a 73.00 –54.47 7.3 –40.4 –2.1 helheim 66.36 –38.12 5.8 –34.6 –1.8 inngia 72.03 –52.61 4.0 –31.7 –1.7 kong oscars 75.98 –59.79 4.2 –21.5 –1.1 sermeq silardleq 70.80 –50.80 3.3 –19.7 –1.0 academy 81.50 –32.65 8.8 –14.2 –0.7 umiamako 71.72 –52.44 2.9 –13.3 –0.7 ikertivaq a 65.67 –39.60 3.2 –13.1 –0.7 docker–smith 76.24 –61.00 5.1 –13.1 –0.7 upernavik b 72.94 –54.38 3.8 –12.4 –0.7 tingmjarmiut 62.76 –43.18 2.5 –11.3 –0.6 ikertivaq d 65.49 –40.06 7.9 –10.6 –0.6 daugaard–jensen 71.92 –28.57 5.3 –10.1 –0.5 sermeq avannarleq 69.36 –50.31 4.4 –9.5 –0.5 hayes 74.92 –57.00 9.6 –9.0 –0.5 perdlerfiup sermia 70.99 –50.92 2.7 –9.0 –0.5 ikertivaq c 65.58 –39.96 5.3 –7.3 –0.4 fenris 66.36 –37.54 2.8 –7.1 –0.4 sermilik 61.00 –45.95 1.5 –6.3 –0.3 kangia nunata sermia 63.33 –49.62 7.8 –6.2 –0.3 ikertivaq b 65.63 –39.64 4.5 –3.6 –0.2 kangerdluarssup sermia 71.25 –51.47 3.2 –3.4 –0.2 upernavik d 72.79 –54.22 2.3 –3.4 –0.2 upernavik e 73.00 –54.65 2.0 –3.4 –0.2 lille 70.43 –50.51 2.1 –2.8 –0.1 upernavik c 72.85 –54.33 6.3 –2.7 –0.1 ryder 81.30 –49.90 8.0 –2.3 –0.1 rink 71.75 –51.64 5.1 2.0 0.1 upernavik f 73.03 –54.84 1.8 –1.9 –0.1 kangigdleq 70.72 –50.64 2.9 –0.8 0.0 qajuuttap 61.32 –45.78 3.2 0.7 0.0 nunatakavsaup sermia 73.22 –55.14 3.6 0.5 0.0 kangerdlugssup sermerssua 71.46 –51.36 4.9 –0.1 0.0 store 70.40 –50.55 5.2 –0.1 0.0 table 2: overview of net area change, average annual area change, and width for the 47 tidewater glaciers monitored between 1999–2018. the table is sorted by magnitude net area change. https://earthexplorer.usgs.gov/ https://earthexplorer.usgs.gov/ https://scihub.copernicus.eu/ e2019430202-03 area change estimates since 1999 visualising annual calving front lines is a useful tool for identifying glacier retreat, advance or stable calving front lines (as shown in fig. 2). comparing sequential annual lines enables annual area change assessment for individual glaciers and for the whole ensemble, by selecting a constant reference line upstream of the glacier. the reference line is somewhat arbitrary, but when combined with each calving front line, we can generate a series of annual polygons. annual area changes are then estimated as the area difference between consecutive polygons. for a thorough presentation of the methodology, see jensen et al. (2016). measuring the polygon area change is arguably a more robust method than measuring a one-dimensional change in glacier length, since it accounts for non-uniform changes to the shape of the calving front. figure 2 shows an example, where the 2012 and 2013 calving front lines of hayes glacier have each been combined with the same reference line, generating two polygons. the 2012–2013 area change is then computed as the 2013 polygon area minus the 2012 polygon area. sources of uncertainty the two main sources of uncertainty in the promice calving line product are the manual delineation of front lines and the timing of front-line mapping. in principle, a front line can be delineated from a given satellite image with the precision of a single image pixel. however, in practice such precise digitisation is not feasible as the exact position of the calving front line is often obstructed by shadows or thin clouds. consequently, an uncertainty of about two image pixels is inherent in the manual digitisation process. on the ground, one pixel represents either 10 m or 15–30 m for the utilised sentinel-2 and landsat bands, respectively. for example, the error associated with the petermann glacier area change is estimated to be around 1 km2 in a 15 m resolution image, or 0.7 km2 using a 10 m resolution image. this rough estimate is computed as four times the image pixel size multiplied by the glacier width (which is about 17.4 km, cf. table 2), representing an uncertainty of ±2 image pixels for each point in the front-line delineation process. the end of the melt season for a given glacier is defined as the time at which the glacier is at its minimum position, i.e. when it has retreated the farthest. however, satellite coverage is limited, and a subset of images are unusable due to the presence of clouds, potentially causing the timing of the end of the melt season to be overor under-estimated. such a mistiming introduces uncertainty when computing annual area change, as area measurements are compared at different points in the season. the final uncertainty depends strongly on how many images are available, and whether the image is obstructed by clouds. jensen et al. (2016) estimated the error due to mistiming to be within 1 km2 and highlighted image availability as the dominant source of error. the problem is somewhat alleviated by the availability of data from sentinel-2, which provides images more frequently due to its short repeat pass time of five days. another source of error arises from an instrument failure on landsat 7’s enhanced thematic mapper plus (etm+) sensor, resulting in bands of missing data across all landsat 7 images from 2003 onwards. in images where one of these bands aligned with the calving front line in parallel, obstructing a major part of the front line, other images from within a few weeks were used to approximate the front-line glacier name lat. long. width net area average area (°n) (°e) (km) change change (km2) (km2 a-1) zachariae 78.90 –20.14 24.6 –409.6 –21.6 petermann 80.10 –61.17 17.4 –289.5 –15.2 humboldt 79.50 –64.61 89.0 –259.4 –13.7 hagen 81.53 –28.50 9.1 –172.8 –9.1 jakobshavn 69.18 –49.73 11.4 –137.0 –7.2 storstrømmen 76.71 –22.47 31.9 –99.5 –5.2 nunatakassaap sermia 74.62 –56.34 5.4 –70.5 –3.7 steensby 81.20 –53.90 4.5 –69.5 –3.7 ostenfeld 81.60 –45.20 7.0 –65.0 –3.4 79 fjorden 79.60 –20.17 42.2 –51.7 –2.7 steenstrup 75.28 –57.89 16.2 –50.8 –2.7 kangerdlugssuaq 68.61 –32.93 6.0 –45.9 –2.4 midgaard 66.45 –36.73 3.8 –40.9 –2.2 upernavik a 73.00 –54.47 7.3 –40.4 –2.1 helheim 66.36 –38.12 5.8 –34.6 –1.8 inngia 72.03 –52.61 4.0 –31.7 –1.7 kong oscars 75.98 –59.79 4.2 –21.5 –1.1 sermeq silardleq 70.80 –50.80 3.3 –19.7 –1.0 academy 81.50 –32.65 8.8 –14.2 –0.7 umiamako 71.72 –52.44 2.9 –13.3 –0.7 ikertivaq a 65.67 –39.60 3.2 –13.1 –0.7 docker–smith 76.24 –61.00 5.1 –13.1 –0.7 upernavik b 72.94 –54.38 3.8 –12.4 –0.7 tingmjarmiut 62.76 –43.18 2.5 –11.3 –0.6 ikertivaq d 65.49 –40.06 7.9 –10.6 –0.6 daugaard–jensen 71.92 –28.57 5.3 –10.1 –0.5 sermeq avannarleq 69.36 –50.31 4.4 –9.5 –0.5 hayes 74.92 –57.00 9.6 –9.0 –0.5 perdlerfiup sermia 70.99 –50.92 2.7 –9.0 –0.5 ikertivaq c 65.58 –39.96 5.3 –7.3 –0.4 fenris 66.36 –37.54 2.8 –7.1 –0.4 sermilik 61.00 –45.95 1.5 –6.3 –0.3 kangia nunata sermia 63.33 –49.62 7.8 –6.2 –0.3 ikertivaq b 65.63 –39.64 4.5 –3.6 –0.2 kangerdluarssup sermia 71.25 –51.47 3.2 –3.4 –0.2 upernavik d 72.79 –54.22 2.3 –3.4 –0.2 upernavik e 73.00 –54.65 2.0 –3.4 –0.2 lille 70.43 –50.51 2.1 –2.8 –0.1 upernavik c 72.85 –54.33 6.3 –2.7 –0.1 ryder 81.30 –49.90 8.0 –2.3 –0.1 rink 71.75 –51.64 5.1 2.0 0.1 upernavik f 73.03 –54.84 1.8 –1.9 –0.1 kangigdleq 70.72 –50.64 2.9 –0.8 0.0 qajuuttap 61.32 –45.78 3.2 0.7 0.0 nunatakavsaup sermia 73.22 –55.14 3.6 0.5 0.0 kangerdlugssup sermerssua 71.46 –51.36 4.9 –0.1 0.0 store 70.40 –50.55 5.2 –0.1 0.0 table 2: overview of net area change, average annual area change, and width for the 47 tidewater glaciers monitored between 1999–2018. 57°15’w 74°57’n 74°55’n 8 km4 620 57°15’w 57°15’w reference line 2013 calving front line 2012 calving front line 2013 polygon 2012 polygon fig. 2: example image from hayes glacier showing how area change is estimated from two consecutive calving front lines. the green line indicates the reference line. the calving front lines in 2012 and 2013 are depicted by blue and red lines, respectively. the glacier area change is computed as the area difference between the 2012 and 2013 polygons. the satellite image is an eu copernicus sentinel-2a image from 14 august 2018. e2019430202-04 position. if a band of missing data crossed the calving front line in a roughly perpendicular way, only obstructing a short stretch of the calving front line, the front line was simply digitised in a straight line across the missing data band. again, this problem is alleviated by the use of sentinel-2 or landsat 8 data. results and discussion as an example of the database, we present the calving front changes of greenland’s fastest flowing glacier, jakobshavn isbræ (sermeq kujalleq; fig. 3). in 2017 the calving front advanced after more than 10 km of retreat between 1999 and 2016. the 2017 advance and stability into 2018 occurred in the faster flowing southern branch of the glacier. a time series of the cumulative net area change for the period 1999–2018, estimated from the calving front line product of 47 glaciers, shows a net area loss of about 2100 km2 (fig. 4). this equates to an average annual area loss of c. 110 km2 for these 47 glaciers since 1999, which corresponds to an area roughly the size of paris. the rate of area change varies substantially through time. the period 2007–2012 underwent a rapid loss of glacier area, compared to 2013–2018, in which glacier area was relatively stable, associated with a small area change. the year 2017–2018 stands out as the only period with net area gain (+4.1 km2). table 2 provides the net area change as well as the average annual area change for each of the 47 surveyed glaciers. nearly all of the investigated marine terminating outlet glaciers show a negative net area change since 1999, indicating a retreat. only three glaciers (rink, nunatakavsaup sermia and qajuuttap) show a positive net area change. there is substantial variation between the glaciers in the ensemble, but a few of the large glaciers, such as humboldt, petermann, zachariae, and jakobshavn, are responsible for a considerable amount of the total ensemble net area change observed in fig. 4. other glaciers such as store and nunatakavsaup sermia show a net area change very close to zero. the area change of two glaciers is not directly comparable; a wider glacier will tend to show larger area changes than a narrower one. alternatively, to compare 69°14’n 50°15’w 50°0’w 49°45’w 49°30’w 69°7’n year 1999 2000 2001 2002 2003 2004 2005 2006 2007 2008 2009 2010 2011 2012 2013 2014 2015 2016 2017 2018 fig. 3: evolution of the calving front position of jakobshavns isbræ from manually delineated end-of-melt-season calving front lines spanning 1999 to 2018. the background image is from the eu copernicus sentinel-2b satellite from 4 september 2018. it is the same image that was used to delineate the 2018 front line. e2019430202-05 the area change of two glaciers, one could compute an estimate of the one-dimensional movement of each glacier as the obtained area change normalised (i.e. divided) by the glacier width. we anticipate the updated promice front line data set to be a useful addition to statistical investigations of glacier behaviour in greenland. questions remain as to the sensitivity of glacier area change to various climate parameters such as atmospheric forcing, sea-surface temperature, sea-ice concentration and the north atlantic oscillation (nao). jensen et al. (2016) found significant correlations between glacier area change and climate parameters for the period 1999–2013, dependent on geographical region. the reduction in ice-area loss in jakobshavn isbræ in 2017–2018 (fig. 3) coincides with anomalously low melt-season temperatures, but kazendar et al. (2019) pointed to colder ocean waters as the primary cause. bevis et al. (2019) supported the notion of inter-annual variations in surface melting, driven by the nao, as an indicator of variability in greenland mass loss. a physical mechanism linking summer air temperature and the front area of marine terminating outlet glaciers is the so-called hydrofracture, in which increased air temperature increases the supply of meltwater. more meltwater increases the availability and weight of water filling surface depressions. water being denser than ice, adds an additional stress that can disaggregate the ice, especially along fractures (e.g. weertman 1973; van der veen 1998). with the promice calving front line product, the time series of calving front line measurements is extended and five more glaciers are added to the ensemble, which serves to improve the accuracy of such statistical investigations performed by jensen et al. (2016). the calving front line product can also be used to compute ice velocity along a glacier flow line. ice velocity products are typically generated by cross-correlation of synthetic aperture radar (sar) images, and do not distinguish between glaciers and sea ice. knowing the position of the calving front line allows a flow line to be constructed based on the glacier alone, thus eliminating this problem. the calving front line product can thus act as an auxiliary data set in cases where annual movement of calving front lines must be taken into account. conclusions the promice calving front line product provides a useful data set of calving front lines at the end of each melt season since 1999. area change estimates generated from earlier versions of the data set have been used in international climate status reporting, for example in the arctic monitoring and assessment programme (amap 2017), in the state of the climate series (e.g. tedesco et al. 2016) and the arctic report card (tedesco et al. 2018). the updated and extended data set now provides annual end-of-melt-season calving front lines for 47 greenland marine terminating glaciers between 1999 and 2018. the data have been added to the list of promice products, which are publicly available at www. promice.dk. acknowledgements we thank one anonymous reviewer and mette kusk gillespie for constructive reviews. the programme for monitoring of the greenland ice sheet (promice) is funded by the geological survey of denmark and greenland (geus) and the danish ministry of energy, utilities and climate under danish cooperation for environment in the arctic (dancea), and is conducted in collaboration with the national space institute (dtu space) and asiaq (greenland survey). we would also like to thank the eu copernicus program and esa for sentinel-2 data, nasa/usgs for landsat data, and nasa/usgs/jss (japans space systems) for aster products. references amap 2017: snow, water, ice and permafrost in the arctic 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https://doi.org/10.1002/2014gl061836 https://doi.org/10.1007/s40641-018-0107-0 https://doi.org/10.1007/s40641-018-0107-0 https://doi.org/10.1038/ngeo394 https://doi.org/10.1126/science.1121381 https://doi.org/10.1126/science.1121381 https://doi.org/10.1175/2016bamsstateoftheclimate.1 https://doi.org/10.1175/2016bamsstateoftheclimate.1 https://arctic.noaa.gov/report-card/report-card-2018/artmid/7878/articleid/781/greenland-ice-sheet https://arctic.noaa.gov/report-card/report-card-2018/artmid/7878/articleid/781/greenland-ice-sheet https://arctic.noaa.gov/report-card/report-card-2018/artmid/7878/articleid/781/greenland-ice-sheet https://doi.org/10.1126/science.1178176 https://doi.org/10.1007/s40641-017-0084-8 https://doi.org/10.1007/s40641-017-0084-8 https://doi.org/10.1016/s0165-232x(97)00022-0 https://doi.org/10.34194/geusb-201943-02-02 geological survey of denmark and greenland bulletin 26, 2012, 53-56 53 anorthosites in greenland: a possible raw material for aluminium? christian knudsen, jan wanvik and henrik svahnberg the famous swiss-born, norwegian geologist and geochemist victor goldschmidt suggested that anorthosite could be used as a source of aluminium replacing bauxite, and acid leaching of the anorthosite was his innovative idea. anorthosite is a rock type consisting of more than 90% plagioclase which is an acid-soluble, aluminium-rich silicate mineral occurring in basement rocks of both norway and greenland (fig. 1). experiments conducted in norway during the century after goldschmidt’s initial idea showed that it is technically possible to use anorthosite as a raw material in the production of aluminium metal. goldschmidt mapped parts of the large anorthosite massifs along sognefjord in the period 1916–1919. during the second world war, sampling and core drilling were conducted in norway, and an anorthosite mine was opened by norsk hydro where up to 400 men were employed and some 15 000 tonnes of rock were quarried before sabotage ended the work in 1945. there was renewed interest in anorthosite as an alternative raw material for aluminium in norway in the years 1976–1982, but experiments conducted in this period did not lead to an economically viable concept. recent developments at the institute for energy technology in norway have led to the discovery of a more promising process based on nitric acid that can yield additional products such as precipitated calcium carbonate (pcc) for the paper industry, amorphous silica and ammonium nitrate fertiliser. the process can also be used as a sink for co2 by taking co2 from, for example, a power plant and binding it to pcc. solubility as a function of mineral chemistry the mineral plagioclase covers a range of compositions from albite (naalsi3o8) to anorthite (caal2si2o8) and forms a solid solution series. the solubility of plagioclase, and therefore of anorthosite, increases with the calcium content, expressed as the anorthite content or an% in the plagioclase (fig. 2). the higher solubility of the calcium-rich plagioclase makes it more attractive as a source of aluminium, as does the content of aluminium which increases with the anorthite content (calcium content; fig. 2). anorthosite bodies in the inner sognefjord–voss area in western norway have a calcium-rich plagioclase composition (fig. 2) with an anorthite content of 65–78%. the solubility and aluminium content determine the quality of anorthosite as a raw material for aluminium production. to provide an overview of archaean (2550–3800) proterozoic (1750–2000) archaean block nuuk neria fiskenæsset buksefjorden qarliit nunaat innajuattoq qaqortorsuaq akia 250 km nagssugtoqidian fold belt volcanic rocks (50–60) sediments (23–416) sediments (416–542) sediments/metasediments/ volcanics (542–1740) gardar intrusions (1120–1350) reworked archaean (1750–2000) godthåbsfjord qaqujârssuaq tunulik boye sø greenland major anorthosite complex minor anorthosite complex sedimentary basins and intrusions precambrian crystalline basement fig. 1. simplified geological map of greenland showing anorthosite occurrences. ages in million years. © 2012 geus. geological survey of denmark and greenland bulletin 26, 53–56. open access: www.geus.dk/publications/bull 5454 the variation in these parameters and accordingly the value of anorthosite as a potential raw material in aluminium production, the geological survey of denmark and greenland conducted a survey of the compositional variation of anorthosite complexes in greenland. anorthosites in greenland anorthosite rock bodies are found in archaean basement rocks in most parts of greenland and constitute up to 5% of the bedrock in a region. they are easy to recognise in the field because of their very light weathering colour which also makes the rock very useful as a structural marker when mapping in deformed terranes. anorthosite can be divided into several types including ‘archaean calcic anorthosite’ (ashwal 1993). the calcium content in the plagioclase of ‘archaean calcic anorthosite’ is high (75–90% an). this feature distinguishes archaean anorthosite from, for example, the ‘proterozoic (massif) type anorthosite’ with 35–60% an. so far, only the archaean type of anorthosite has been described from greenland. the archaean anorthosite occurrences in greenland are generally deformed and metamorphosed to such an extent that their genetic relationships are difficult to reveal. they often occur as decimetreto metre-sized pods and inclusions in the country gneiss. however, there are a number of places where larger bodies of anorthosites are found with preserved primary textures and relationships. the most prominent occurrence is the fiskenæsset complex consisting of anorthosite, leucogabbro, gabbro and ultramafic rocks. here it has been demonstrated that anorthosite forms parts of large intrusions of basaltic composition and formed as cumulates by crystal fractionation (windley et al. 1973; myers 1975; windley & garde 2009). 100 80 60 40 20 0 g re en la n d 0 20 40 % an of plagioclase 60 80 100 l ea ch ed a l 2 o 3 ( % o f to ta l a 2 o 3 ) 10 30 50 70 90 oligoclase andesine labrabytownite a n o rt h it e doritea lb it e 100 80 60 40 20 0 g re en la n d 0 20 40 % an of plagioclase 60 80 100 l ea ch ed a l 2 o 3 i n % o f to ta l a 2 o 3 10 30 50 70 90 oligoclase andesine labrabytownite a n o rdoritea lb it e naalsi3o8 caal2si2o8 th it e unaltered gudvangen type rogaland type so gn u n al te re d g u d va n ge n t yp esio 2 na 2o al2o3 cao fig. 2. chemical composition (left) and solubility (right) of plagioclase. modified from wanvik (2000). gudvangen, rogaland and sogn are anorthosite occurrences in norway. greenland is based on an average from table 1. table 1. average al 2 o 3 , cao, fe 2 o 3 and na 2 o compositions of anorthosites calculated from whole-rock analytical data fiskenæsset group 1 † 63°15´ 50°00´ 13 30.6 16.3 2.5 1.8 90.1 83.8 fiskenæsset group 2 3 32.5 16.3 1.9 1.8 94.1 83.6 fiskenæsset group 3 3 30.5 12.5 1.5 4.0 92.4 65.7 buksefjorden 63°55´ 51°20´ 6 29.3 14.1 1.5 2.2 90.7 73.7 qarliit nunaat (godthåbsfjord) 64°04´ 49°45´ 1 28.5 11.9 0.9 4.1 93.0 61.0 naajat kuuat (godthåbsfjord) 64°10´ 50°03´ 7 28.9 13.6 1.8 2.5 88.7 72.9 storø (godthåbsfjord) 64°23´ 51°07´ 1 32.3 16.0 0.6 2.1 96.6 81.3 akia (godthåbsfjord) 64°30´ 52°06´ 9 30.6 15.2 1.8 2.2 91.4 79.6 ivisaartoq (godthåbsfjord) 64°44´ 49°42´ 6 29.8 14.7 3.0 2.6 91.1 75.2 innajuattoq (godthåbsfjord) 64°45´ 50°40´ 1 31.8 16.0 1.5 2.0 91.6 84.4 qaqortorsuaq 66°35´ 52°12´ 3 33.6 15.8 0.9 2.2 94.1 82.6 tunulik 70°03´ 51°15´ 5 28.9 12.9 1.0 3.5 91.5 67.1 qaqujârssuaq 77°35´ 64°45´ 2 29.9 14.9 1.2 1.7 91.6 77.3 gudvangen (norway) 8 30.1 14.1 0.8 2.9 94.6 72.0 n w no of cipw cipw lat. long. analyses al2o3 cao fe2o3 na2o % plag* % an* * % plagioclase in rock and % an in plagioclase are based on cipw norm calculations. † fiskenæsset is grouped on the basis of its rareearth element patterns following polat et al. (2009). 55 the primary relationship to the surrounding rocks is often obscured by tectonic activity or intrusive contacts to younger granitoids. all the anorthosites studied here are assumed to belong to the calcic archaean type but their composition varies, (1) among the different complexes, (2) within the complexes as a function of the stratigraphical position and (3) within the mineral grain – from core to rim, often due to recrystallisation during metamorphism. around 12 anorthosite complexes in greenland have been mapped and described (fig. 1). the northernmost one is the qaqujârssuaq anorthosite, which is also the largest single anorthosite mass in greenland covering c. 100 km2 of smithson bjerge and an unknown area under the inland ice (dawes 2006). it is a c. 500 m thick succession composed of c. 90% anorthosite, c. 10% leucogabbro and <1% gabbro which was emplaced c. 2700 ma ago (nutman 1984). the tunulik anorthosite is located in an area of archaean rocks deformed and metamorphosed in palaeoproterozoic time c. 1900 ma ago. generally, the anorthosite occurs as blocks and pods in the surrounding tonalitic to granodioritic gneisses (andersen & pulvertaft 1986). the anorthosite can be traced south to the c. 25 km2 large boye sø anorthosite (garde & steenfelt 1989). the first anorthosite body to be found in greenland was the anorthosite at qaqortorsuaq in the c. 1900 ma old palaeoproterozoic nagssugtoqidian fold belt (ellitsgaardrasmussen & mouritsen 1954). this body is very large and the exploration company kryolitselskabet øresund a/s estimated that there are c. 100 million tonnes of anorthosite per vertical metre in the deposit and the mountain qaqortorsuaq is c. 1300 m high (gothenborg & keto 1977). the highest concentration of anorthosite complexes in greenland is found in the core of the archaean block around nuuk (akia, innajuattoq, storø, najaat kuuat with qarliit nunaat, nunatuasuk and ivisaartoq; fig. 1; table 1). one of the anorthosite bodies is located at innajuattoq (fig. 3). the fiskenæsset anorthosite complex is one of the largest and best known archaean anorthosite complexes worldwide. parts of the complex have retained an igneous stratigraphy, cumulate textures, layering, grading and channel deposits (windley et al. 1973; windley & smith 1974; windley & garde 2009; myers 1975, 1976, 1985), showing that it is a sheet-like, layered basic intrusion. based on in a b fig. 3. a: the anorthosite at innajuattoq in the archaean block northwest of nuuk. the mountain is 1206 m high, and the cliff section is c. 1100 m high. b: close-up view; length of hammer handle c. 60 cm. % c ao 6 8 10 12 14 16 18 20 22 qaqortorsuaq akia innajuattoq naajat kuuat (godthåbsfjord) qarliit nunaat fiskenæsset 24 26 28 30 32 34 36 38 % al2o3 fig. 4. microprobe analyses of plagioclase from greenland anorthosites showing cao versus al2o3 wt%. the range of plagioclase compositions is due to variations from core to rim in individual grains as well as to variations between different parts of the complexes. the linear relationship between cao and al2o3 is due to the coupled substitution of albite (naalsi3o8) by anorthite (caal2si2o8). 8 9 10 11 12 13 14 15 16 17 18 20 22 24 26 28 30 32 34 36 fiskenæsset 1 fiskenæsset 2 fiskenæsset 3 fiskenæsset others buksefjorden qarliit nunaat naajat kuuat (godthåbsfjord) storø (godthåbsfjord) innajuattoq akia qaqortorsuaq tunulik qaqujârssuaq gudvangen (norway) % al2o3 % c ao fig. 5. whole-rock analyses of anorthosites showing wt% cao versus wt% al2o3. anorthosites in greenland are compared to anorthosites from gudvangen in norway. most of the greenlandic occurrences have higher cao content due to higher anorthite content. hence their plagioclase is rich in both ca and al. 5656 situ 207pb/206pb zircon ages of up to 2950 ma, keulen et al. (2010) concluded that the intrusion age is c. 2970 to 2950 ma. the anorthosite unit is c. 250 m thick (myers 1985). at localities where the anorthosite is least deformed, it typically appears as megacrystic with 1–10 cm equant, relict igneous plagioclase grains dispersed in 1–5 mm large metamorphic plagioclase. the main part of the anorthosite is deformed, and the plagioclase is metamorphic in a granular texture. anorthosite composition there is a large range of plagioclase compositions among the anorthosite complexes in greenland (fig. 4) with the highest content of calcium found in the fiskenæsset complex. whole-rock analyses indicate that very calcium-rich plagioclase also occurs in the anorthosite at akia (fig. 5; dymek & owens 2001). most greenland anorthosite rocks are more calcic than the norwegian ones and according to the norwegian experiences should be more soluble and hence more suitable as a raw material for aluminium production. a possible continuation of this project could be to collect samples from anorthosites in greenland and conduct solubility tests using the norwegian methods. the most promising occurrence is the fiskenæsset complex which is the largest anorthosite in greenland, has the highest bulk rock cao content and contains the most calcic plagioclase in greenland. the anorthosites at akia, innajuattoq and qaqortorsuaq contain very calcic plagioclase and low contents of other minerals. these occurrences are located close to the sea and could be targeted in further studies. acknowledgement the bureau of minerals and petroleum in nuuk is thanked for financial support. references andersen, m.c. & pulvertaft, t.c.r. 1986: occurrences of anorthositic rocks in reworked archaean basement in the umanaq area, central west greenland. rapport grønlands geologiske undersøgelse 129, 18 pp. ashwal, l.d. 1993: anorthosites, 422 pp. berlin: springer. authors’ addresses c.k., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ckn@geus.dk. j.w., geological survey of norway, leiv eirikssons vei 39, 7040 trondheim, norway. h.s., stockholm university, svante arrhenius väg 8, se-106 91, stockholm, sweden. dawes, p.r. 2006: explanatory notes to the geological map of greenland, 1:500 000, thule, sheet 5. geological survey of denmark and greenland map series 2, 97 pp. dymek, r.f. & owens, b.e. 2001: chemical assembly of archaean anorthosites from amphiboliteand granulite-facies terranes, west greenland. contributions to mineralogy and petrology 141, 513–528. ellitsgaard-rasmussen, k. & mouritzen, m. 1954: an anorthosite occurrence from west greenland. meddelelser fra dansk geologisk forening 12, 436–442. garde, a.a. & steenfelt, a. 1989: a new anorthosite/gabbro complex at nûgssuaq, central west greenland. rapport grønlands geologiske undersøgelse 145, 16–20. gothenborg, j. & keto, l. 1977: report on the aerial reconnaissance between sukkertoppen ice calot and nordenskiölds gletscher 1977, 84 pp. unpublished report, kryolitselskabet øresund a/s, copenhagen, denmark (in archives of geological survey of denmark and greenland, geus report file 20210). keulen, n., næraa, t., kokfelt, t.f., schumacher j.c. & scherstén a. 2010: zircon record of the igneous and metamorphic history of the fiskenæsset anorthosite complex in southern west greenland. geological survey of denmark and greenland bulletin 20, 67–70. myers, j.s. 1975: igneous stratigraphy of archaean anorthosite at majorqap qâva, near fiskenæsset, south-west greenland. rapport grønlands geologiske undersøgelse 74, 27 pp. myers, j.s. 1976: channel deposits of peridotite, gabbro and chromitite from turbidity currents in the stratiform fiskenæsset anorthosite complex, southwest greenland. lithos 9, 281–291. myers, j.s. 1985: stratigraphy and structure of the fiskenæsset complex, southern west greenland. bulletin grønlands geologiske undersøgelse 150, 72 pp. nutman, a. 1984: precambrian gneisses and intrusive anorthosite of smithson bjerge, thule district, north-west greenland. rapport grønlands geologiske undersøgelse 119, 31 pp. polat, a., appel, p.w.u., freyer, b., windley, b., frei, r., samson, i.m. & huang, h. 2009: trace element systematics of the neoarchaean fiskenæsset anorthosite complex and associated meta-volcanic rocks, sw greenland: evidence for a magmatic arc origin. precambrian research 175, 87–115. wanvik, j.e. 2000: norwegian anorthosites and their industrial uses, with emphasis on the massifs of the inner sogn–voss area in western norway. norges geologiske undersøkelse bulletin 436, 103–112. windley, b.f. & garde, a.a. 2009: arc-generated blocks with crustal sections in the north atlantic craton of west greenland: crustal growth in the archean with modern analogues. earth-science reviews 93, 1–30. windley, b.f. & smith, j.v. 1974: the fiskenæsset complex, west greenland. part ii. general mineral chemistry from qeqertarssuatsiaq. bulletin grønlands geologiske undersøgelse 108, 54 pp. windley, b.f., herd, r.k. & bowden, a.a. 1973: the fiskenæsset complex, west greenland. part i. a preliminary study of the stratigraphy, petrology and whole rock chemistry from qeqertarssuatsiaq. bulletin grønlands geologiske undersøgelse 106, 80 pp. geological survey of denmark and greenland bulletin 38, 2017, 13-16 13 buried valleys are elongate erosional structures in the danish subsurface now partly or completely filled and covered with younger sediments. the majority was formed by meltwater underneath ice sheets. the number of buried-valley structures in denmark is large, and because the valley-infill in many areas hosts significant groundwater resources, knowledge of them and their formation is important. this was the starting point of the buried-valley mapping project, which was initiated in the late 1990s and continued until the end of 2015 (sandersen & jørgensen 2016). this project became part of the national groundwater mapping programme which was set up with the purpose of mapping the groundwater resources within areas of specific groundwater interest (thomsen et al. 2004). the areas of specific groundwater interest encompass existing catchment areas and cover around 40% of the country. within these areas, high-density electromagnetic surveys have typically been performed together with exploration drilling and supplementary geophysical measurements. the mapping of the buried valleys has been based on these newly collected data as well as existing data in the national databases. in some instances, it has also been possible to map buried valleys in less data-dense areas outside the surveyed areas, mainly on the basis of borehole data. the groundwater resource and its vulnerability have been important in the mapping of the buried valleys. the valleys also constitute an important part of the subsurface geological architecture, and it is obvious that a thorough knowledge of them is critical for the general understandburied tunnel valleys in denmark and their impact on the geological architecture of the subsurface peter b.e. sandersen and flemming jørgensen buried valley surveyed area 50 km fig. 1. mapped buried valleys in denmark as of end 2015. the mapped valleys are shown as dark grey polygons and the tem-surveyed areas are shown in light grey. the haderslev area covered by fig. 2 is highlighted with a red rectangle. © 2017 geus. geological survey of denmark and greenland bulletin 38, 13–16. open access: www.geus.dk/publications/bull 1414 ing of the geology of the uppermost 100–400 m of the danish subsurface. in this paper we present an overview of the buried-valley mapping project and an updated buriedvalley map (fig. 1). mapping of the buried-valley structures all relevant geophysical and lithological data from primarily onshore areas have been used to map and describe the buried valleys. the transient electromagnetic method (tem; sørensen & auken 2004) has proven especially valuable because such surveys usually provide a spatially dense data grid that can be combined with borehole data and other geophysical data (jørgensen & sandersen 2009). the valleys have been delineated from an integrated interpretation of the data, and their outlines and extensions within the mapped areas were drawn as polygons using simple signatures (fig. 2). in order to obtain a high degree of certainty and objectivity in the delineation of the valleys, the lateral extent and orientation of the valleys were to be unambiguously expressed in the data. therefore, no interpolations outside and between the local mapped areas have been made, and accordingly the map in fig. 1 represents the minimum occurrence of buried valleys in denmark rather than their true distribution and density. the advantage of this approach is an un-biased picture of the valleys based on the data available and not on secondary data and assumptions. in addition to this, the approach gives an opportunity to get a valuable insight into the formation history and age of the valleys. the highest valley densities are found in areas where tem data have been collected and the conditions for the chosen methods were ideal. the map in fig. 1 only shows the location of the valley structures where they can be outlined by interpretation of the data. the grey areas indicate the areas where tem data are available and many additional buried valleys must be expected outside these areas. although many buried valleys have been mapped, even more are expected to exist, because not all valleys can be identified with the used methods. for instance, very narrow valleys or channels can be difficult to resolve, and valleys with low lithological and/or electrical resistivity contrast compared to the surroundings can be difficult to outline. occurrence and subsurface architecture the depth of the mapped buried valleys is variable, with the deepest structures sometimes exceeding 400 m. their width is generally between 0.5 and 1.5 km, but widths of more than 3.5 km occur. their lengths are difficult to assess because many of the surveyed areas are small, but some exceed 25–30 km. they commonly terminate abruptly and are highly irregular, with depressions and thresholds along the valley floors. buried valleys appear both as single valleys and in dense cross-cutting networks. their internal structure is typically complex due to repeated erosional and depositional events. the majority of the buried valleys were formed during the pleistocene as tunnel valleys eroded by high-pressure meltwater underneath the ice sheets (jørgensen & sandersen 2006). with respect to morphology and dimensions, a b c d e rib62 f partially buried (open) valleys poorly documented in data well documented in data completely buried valleys poorly documented in data well documented in data 5 km halk hoved a b 5 km 2000 100 tem mean resistivity (ohm-m) fig. 2. survey area rib62 haderslev, south-western denmark (outlined by blue polygon). a: airborne tem resistivity data (30–35 m b.s.l.). b: same map window without tem data but with interpreted buried valleys shown as hatched polygons. green line shows location of cross-section in fig. 3. tunnel valleys marked a to f corresponds to valleys shown in fig. 3. 15 the mapped buried valleys are comparable with open tunnel valleys found in the present-day danish landscape. the water seemed to flow in relatively small channels on the floors of the tunnel valleys, which gradually became icefilled. tunnel valleys were often re-used during repeated cycles of glaciations, producing separate valley generations and multiple internal cut-and-fill structures. the ages of individual valleys are usually difficult to assess because of the repeated erosion and the general lack of precise age determinations, but it is assumed that tunnel-valley formation has been a common phenomenon throughout the quaternary. relative dating of the individual generations can commonly be performed, based on the lithology of the infill and on cross-cutting relationships in the subsurface. in rare cases mapped valley generations can be related to specific ice advances (i.e. sandersen et al. 2009). even when taking the irregular distribution of the survey areas into consideration, the mapped valleys show signs of a preferred geographical distribution (fig. 1). for instance, the highest densities of buried tunnel valleys are typically found in areas where thick successions of palaeogene clays occur close to the surface, and the lowest densities typically where coarse-grained sediments dominate the near-surface part of the succession. the tunnel-valley formation is thus more likely to occur in areas with impermeable substrata because the high subglacial water pressures favour channelised erosion and tunnel-valley formation. however, in areas where the permeable sediments underneath the glacier were able to drain parts of the meltwater, no or only few valleys were formed. apparently, an important factor is whether drainage through the substrata can be sufficient to prevent tunnel-valley formation (sandersen & jørgensen 2012). a map of the pre-quaternary surface in denmark by binzer & stockmarr (1994) reveals a deep valley-network eroded into the pre-quaternary surface. this map was primarily based on on-shore borehole data. the interpolation of the pre-quaternary surface was highly interpretative and performed throughout the danish area under the presumption that the structures were sub-aerially eroded valleys and would therefore not contain isolated depressions and blind endings. it is tempting to compare the map by binzer & stockmarr with the map in fig. 1, because some of the valley structures coincide. however, the buried-valley map in fig. 1 is made differently and is not a map of the pre-quaternary surface. the map shows the occurrence of buried-valley structures regardless of whether they penetrate the pre-quaternary surface or not and hence the map is not limited to showing valleys in the pre-quaternary surface, and isolated depressions and blind endings have not been avoided when making the map. a total length of 5600 km of buried valleys has been found within the tem-mapped areas in denmark (c. 17 000 km2) and in selected adjoining areas (c. 1700 km2); in total c. 18 700 km2. if the average width of the valleys is 1 km (see jørgensen & sandersen 2006) the valleys cover 5600 km2. this means that in c. 30 % of the mapped area, erosion of the valleys has significantly changed the lithology and architecture of the subsurface. an example where mapped buried tunnel valleys dominate c. one third of the area is illustrated in fig. 2. the tem data show that tunnel valleys in this case have been eroded through the quaternary and miocene deposits and in some instances deeply into the underlying palaeogene clays (fig. 3). the cross-section in fig. 3 illustrates the influence the buried tunnel valleys have on the subsurface architecture. the pre-quaternary succession of miocene sands and clays have been removed and replaced with sandy and clayey infill of quaternary age. figure 2a shows a horizontal slice (30–35 m b.s.l.) approximately through the central part of the miocene succession. the red to orange colours show high resistivities corresponding to predominantly sandy sediments. the buried tunnel valleys are seen as both green elongate structures where the green colours represent predominantly clayey sediments and red elongate structures consisting of predominantly sand. as seen on the cross-sec    miocene  miocenemiocene? palaeogene palaeogene   quaternary quaternary a b c d e f 0 5 10 15 19 km -200 -150 -100 0 50 m a.s.l. n s fig. 3. cross-section of tunnel valleys a to f in survey area rib62 haderslev based on 3d-gridded airborne tem resistivity data. vertical exaggeration 15 times. vertical rods are boreholes included in the national jupiter database. see fig. 2 for resistivity legend and location. 1616 tion (fig. 3) valley erosions reach down to the palaeogene clays, seen as very low resistivities in the tem data (blue colours). the valleys show dominant orientations of ne–sw and se–nw, respectively (fig. 2b). the valley infill is of quaternary age, but the precise age of the valleys is not known. however, relative ages of the individual valley generations can be inferred from cross-cutting relationships, the terrain and the character of the overlying sediments. the buried tunnel valley d in fig. 2b, for example, is partially buried and is expected to extend all the way to the terrain surface because it coincides with an open tunnel valley in the present-day terrain (smed 1982). the valley obviously belongs to a young generation. the deep valley c, however, is completely buried by predominantly clay tills and apparently not reaching the surface. this valley is therefore most likely older than the adjacent valley d to the south-east. the valley e and its extension to the north-west appear to be cut by the valleys c, d and f, thus suggesting an older generation. the southern part of valley e is quite easily seen in the tem-data from c. 30 to 100 m b.s.l., whereas the levels above the succession are dominated by low-resistivity clays apparently not belonging to the valley. at the nearby coastal cliff of halk hoved in the southeasternmost part of the area (fig. 2), a large glaciotectonic complex formed by proglacial deformation of the ne-advance in late weichselian has been described by madsen & piotrowski (2012). this thrust-fault complex consists of stacked layers of tills and glaciofluvial sediments deposited during the warthe glaciation in late saalian. according to madsen & piotrowski (2012), the decollement layer of the glaciotectonic complex is c. 20 m b.s.l. which is apparently above the level of the buried valley e farther to the south-west. the uppermost parts of the buried tunnel valley may have been deformed by the glaciotectonic event in late weichselian, but obviously the formation of the deep tunnel valley can be related to an earlier event. conclusions and perspectives if the rough calculation above is extrapolated to cover all of the danish onshore area, a plausible total length of buried valleys of around 13 000 km emerges. despite the uncertainty of this calculation, the figure calls for attention when interpreting the subsurface geology of areas not at present covered by dense geophysical datasets. mapping outside the areas with specific groundwater interests is at present sparse, but the general knowledge obtained from the mapping project can be used in areas not yet covered by dense datasets. along with the continuously increasing knowledge of the occurrence and the origin of buried tunnel valleys in denmark, the importance of their structures has become more and more evident. the importance for assessments of groundwater resources and their vulnerability is straightforward (i.e. andersen et al. 2013; sandersen & jørgensen 2003), and the necessity of making the buried-valley structures play a significant role in the geological interpretations of quaternary sedimentary sequences is unquestionable. references andersen, t.r., poulsen, s.e., christensen, s. & jørgensen, f. 2013: a synthetic study of geophysics-based modelling of groundwater flow in catchments with a buried valley. hydrogeology journal 21, 491–503. binzer, k. & stockmarr, j. 1994: geologisk kort over danmark. prækvartæroverfladens højdeforhold. danmarks geologiske undersøgelse kortserie 44, 10 pp. jørgensen, f. & sandersen, p. 2006: buried and open tunnel valleys in denmark – erosion beneath multiple ice sheets. quaternary science reviews 25, 1339–1363. jørgensen, f. & sandersen. p.b.e. 2009: buried valley mapping in denmark: evaluating mapping method constraints and the importance of data density. zeitschrift der deutschen gesellschaft für geowissenschaften 160, 211–223. madsen, t.m. & piotrowski, j.a. 2012: genesis of the glaciotectonic thrust-fault complex at halk hoved, southern denmark. bulletin of the geological society of denmark 60, 61–80. sandersen, p. & jørgensen, f. 2003: buried quaternary valleys in western denmark – occurrence and inferred implications for groundwater resources and vulnerability. journal of applied geophysics 53, 229–249. sandersen, p.b.e., jørgensen, f. larsen, n.k., westergaard, j.h. & auken, e. 2009: rapid tunnel-valley formation beneath the receding late weichselian ice sheet in vendsyssel, denmark. boreas 38, 834–851. sandersen, p.b.e. & jørgensen, f. 2012: substratum control on tunnelvalley formation in denmark. in: huuse, m. et al. (eds) 2012: glaciogenic reservoirs and hydrocarbon systems. geological society special publications (london) 368, 145–157, http://dx.doi.org/10.1144/ sp368.12 sandersen, p.b.e. & jørgensen, f. 2016: kortlægning af begravede dale i danmark. opdatering 2010–2015. vols 1, 106 pp. and 2, 626 pp. in danish. geological survey of denmark and greenland, special publication. available from www.buried-valleys.dk smed, p. 1982: landskabskort over danmark, blad 3, sønderjylland, fyn. geografforlaget, brenderup. sørensen, k.i. & auken, e. 2004: skytem. a new high-resolution helicopter transient electromagnetic system: exploration geophysics 35, 194–202, http://dx.doi.org/10.1071/eg04194 thomsen, r., søndergaard, v.h. & sørensen, k.i. 2004: hydrogeological mapping as a basis for establishing site-specific groundwater protection zones in denmark. hydrogeology journal 12, 550–562. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: psa@geus.dk geological survey of denmark and greenland bulletin 1, 585-609 585 the lower–middle jurassic of the anholt borehole: implications for the geological evolution of the eastern margin of the danish basin ole b. nielsen, marit-solveig seidenkrantz, niels abrahamsen, birthe j. schmidt, eva b. koppelhus, helle ravn-sørensen, uffe korsbech and k. gynther nielsen this study of upper pliensbachian – bajocian/bathonian deposits in a borehole drilled on the island of anholt, denmark incorporates sedimentology, biostratigraphy (palynomorphs and foraminifera), palaeomagnetism and coal petrology. the studied succession records a gradual change from marine inner shelf storm-influenced clays to mainly terrestrial sands, clays, and lignite containing a flora of mainly freshwater algae and pollen. the regression was initiated at the pliensbachian–toarcian boundary and marine influence ceased during bajocian–bathonian times; the regression thus took place earlier at anholt than in the centre of the danish basin. the sediments in the anholt borehole are referred to the fjerritslev and haldager sand formations. although the lower–middle jurassic boundary is commonly placed at the boundary between the two formations, our data indicate that at anholt the upper fjerritslev formation (member f-iv) is of aalenian age. the lower–middle jurassic boundary occurs close to the boundary between members f-iii and f-iv of the fjerritslev formation. in contrast to other lower–middle jurassic successions in the north sea region, smectites of inferred volcanic origin are preserved in the anholt section, suggesting limited burial and hence less intense diagenetic illitisation or chloritisation of smectites. a down-hole increase in diagenetic influence is reflected by the increase down-section both in the thermal stability of kaolinite and in the vitrinite reflectance. kaolinite of inferred authigenic origin forms a white powder in the quartz-dominated sands of the haldager sand formation; this kaolinite is thermally very unstable and is interpreted to be of late diagenetic, post-uplift origin. the vitrinite reflectance data indicate that the jurassic formations have been exposed to thermal maturation corresponding to burial to a depth of 1000–1200 m below their present depth. post-maturation uplift of the order of 1 km probably occurred partly during late cretaceous – paleocene inversion in the kattegat area and partly during oligocene–recent regional uplift, the latter being the most important of the two uplift phases. palaeomagnetic data indicate that the main carrier of magnetic remanence is fine-grained magnetite. the stable remanence shows a pronounced inclination shallowing, which is attributed to post-depositional compaction. keywords: danish basin, fennoscandian border zone, lower–middle jurassic, anholt borehole, biostratigraphy, palaeomagnetics, sedimentology, clay mineralogy, organic petrology, geochemistry o.b.n. & m.-s.s., department of earth sciences, university of aarhus, c.f. møllers allé, dk-8000 århus c, denmark. e-mail: geololen@aau.dk n.a., department of earth sciences, university of aarhus, finlandsgade 6–8, dk-8200 århus n, denmark. b.j.s., statoil, exploration division, forushagen, n-4035 stavanger, norway. e.b.k., geological survey of denmark and greenland, geocenter copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. present address: royal tyrrell museum of palaeontology, box 7500, drumheller t0j 0y0, alberta, canada. h.r.-s., samfundsteknik a/s, vesterballevej 4–6, fredericia miljøcenter, dk-7000 fredericia, denmark. present address: bascon a/s, åboulevarden 21, p.o. box 510, dk-8100 århus c, denmark. u.k. & k.g.n.*, ørsted-dtu, building 327, technical university of denmark, dk-2800 lyngby, denmark. *retired. geological survey of denmark and greenland bulletin 1, 585–609 (2003) © geus, 2003 586 during early and middle jurassic times, a major depositional basin, the danish basin, extended across most of denmark and into the modern north sea. it was demarcated by the baltic shield to the north and northeast and by the ringkøbing–fyn high to the south, and was linked with the danish–polish trough to the southeast (sorgenfrei & buch 1964; michelsen 1978; liboriussen et al. 1987). previous studies have shown that the upper part of the lower jurassic fjerritslev formation was characterised by a transgressive phase followed by a regressive phase, from open marine conditions during the pliensbachian – early toarcian to a more restricted marine setting during the late toarcian (michelsen 1978, 1989a). during the middle jurassic, the prograding deltaic or braided river sediments of the haldager sand formation were deposited in the northeastern part of the basin (michelsen 1978, 1989b; koch 1983; michelsen et al. 2003, this volume). this study presents a lower–middle jurassic section from a 306 m deep borehole, drilled in 1990 on anholt, a small danish island in the centre of the kattegat (fig. 1). during the jurassic, anholt was situated near the eastern margin of the danish basin. a composite section was attained by means of three boreholes, anholt ii, iii, and iv, spaced less than 20 m apart. only the latter two extended into jurassic sediments. the boreholes penetrated 104 m of quaternary sediments before reaching the jurassic. the drilling site has an elevation of about 2 m above present-day sea level. all depths noted are referred to this point. the aims of the study were to evaluate the age, the palaeoenvironmental setting and the post-depositional history of the jurassic succession. to achieve this goal, studies were carried out in biostratigraphy, sedimentology, coal petrography, and palaeomagnetism. the material from 104–230 m consists largely of ditch cutkattegat sweden anholt jylland denmark uglev-1 frederikshavn city-1 skagen-3 50 km 2 km 8°e 57°n 55°n 12°e terne-1 gassum-1hobro-1 stenlille wells bornholm skåne skagerrak ringkøbing– fyn high danish basin skagerrak–kattegat platform sorgenfrei–tornquist zone dpt øresund well normal fault areas of late cretaceous and early tertiary inversion basement high børglum-1 fig. 1. map showing the structural outline of the fennoscandian border zone (sorgenfrei–tornquist zone and skagerrak–kattegat platform) and the danish basin, and the location of boreholes and outcrops mentioned in the text (structural outline from petersen et al. 2003, this volume). the inset map of anholt shows the location of the investigated borehole. dpt, danish–polish trough. 587 tings, together with the few available core samples, whereas the samples from below 230 m are all core samples (fig. 2). below 250 m, core recovery was almost 100%, whereas the interval above 250 m is only represented by few, incomplete cores. the latter cores seem mainly to represent clay or silt-rich lithologies, whereas levels with sand are less commonly represented, and primary sedimentary structures might have been destroyed. in ditch cuttings samples, the proportion of clay-sized particles is probably under-represented, as part of the only slightly consolidated clay-rich layers was incorporated with the drilling fluid during coring. 11.0 48.3 9.5 14.1 24.9 150 200 250 300 1 3 5 7 0 40 80 li th os tr at ig ra ph y to ta l o rg an ic ca rb on ( to c ) (w t% ) r ew or ke d pa ly no m or ph s (% ) pa ly no m or ph zo ne s fo ra m in ife ra l zo ne s d ep th ( m b .s .) h al da ge r sa nd f or m at io n fj er ri ts le v fo rm at io n fiii m em be r fiv m em be r pd ba rr en m id dl e ju ra ss ic lo w er ju ra ss ic pc fd fc fb pb fapa quat. c hr on os tr at ig ra ph yfig. 2. chronostratigraphy, biostratigraphical zonation, total organic carbon (toc) content, and reworked palynomorphs of the lower 200 m of the anholt borehole (slightly modified from seidenkrantz et al. 1993). quat., quaternary; m b.s., metres below surface. biostratigraphy the age of the sediments is based on the biostratigraphical analysis. this part of the study has been previously reported in detail (seidenkrantz et al. 1993) and here we only present a synopsis of that study. the preparation of the samples was undertaken using standard techniques (meldgaard & knudsen 1979; dybkjær 1988; poulsen et al. 1990) and the analysis resulted in the establishment of four palynological and four benthic foraminiferal zones (fig. 2). microfauna foraminifera are only present in the lower part of the section and cannot be correlated with faunas from other deposits in the region. with the exception of the lowermost zone (fa), the assemblages almost exclusively consist of agglutinated species (seidenkrantz et al. 1993), whereas jurassic foraminiferal assemblages in surrounding areas are dominated by calcareous species, mainly nodosarids (nørvang 1957; bang 1968a, b, 1971, 1973; norling 1972). in general, the foraminiferal assemblages have a very low diversity and several samples are barren. the lower foraminiferal zone (fa) mainly contains nodosarid genera such as astacolus and planularia. the deposit also contains ammonite fragments, bivalves, and gastropods. in contrast, zone fb (304–288 m) is dominated by the agglutinated genera ammobaculites, bulbobaculites, kutsevella, and haplophragmoides. the fauna of zone fc is essentially restricted to one species of ammobaculites, whereas zone fd yields ammobaculites, bulbobaculites and a few kutsevella (seidenkrantz et al. 1993). the foraminifera are of a boreal affinity not previously found this far to the south (nagy & seidenkrantz in press). palynology the palynomorph assemblages allow close correlation to other sections from denmark, germany, and britain (fig. 3; schulz 1967; lund 1977; bertelsen 1979; woollam & riding 1983; hoelstad 1985; dybkjær 1991; riding & thomas 1992; koppelhus & nielsen 1994; koppelhus & batten 1996). palynomorph zone pa based on the presence of nannoceratopsis spp., luehnda spinosa morgenroth 1970, and mendicodium reticulatum morgenroth 1970. this assemblage correlates well with zone c3 in the hobro-1 borehole in central jylland (bertelsen 1979) and the top of the cerebrocellenites macrooverrucosus zone in the gassum-1 borehole in central jylland (dybkjær 1991) from the danish basin and with the luehndea spinosa zone in the british part of the north sea (riding & thomas 1992). further possible correlations are with the luehnda spinosa zone and the uppermost chasmatosporites zone, both from the bagå formation of bornholm (koppelhus & nielsen 1994) in the danish–polish trough and from the rya formation in the øresund area (fig. 3; koppelhus & batten 1996). recent palynological study of the neill klinter group, jameson land, east greenland (koppelhus & dam 2003, this volume) permits correlation of zone pa from anholt with assemblage zone 2 of the neill klinter group at the albuen section. the combined evidence suggests a late pliensbachian age for zone pa. palynomorph zone pb based on the acme of spheripollenites together with corollina, ischyosporites variegatus (couper) schulz 1967 and manumia delcourtii (pocock) dybkjær 1991. the zone is correlated with zone c4 from the hobro-1 borehole (bertelsen 1979) and the spheripollenites– leptolepidites zone in the stenlille boreholes (dybkjær 1991) of the danish basin. it can also be correlated with zone i of the bagå formation (hasle klinkerfabrik clay pit of bornholm; hoelstad 1985) and the nannoceratopsis gracilis and spheripollenites–leptolepidites zones from the bagå formation (korsodde section, bornholm; koppelhus & nielsen 1994) in the danish– polish trough. the spheripollenites–leptolepidites zone is also known from the rya formation in the øresund area (koppelhus & batten 1996) and from assemblage zone 5 in the neill klinter group in east greenland (koppelhus & dam 2003, this volume). the age is considered to be toarcian. palynomorph zone pc based on abundant perinopollenites elatoides couper 1958 together with callialasporites turbatus (balme) 588 589 m id dl e ju ra ss ic ba th on ia n ba jo ci an a al en ia n to ar ci an lo w er ju ra ss ic se ri es st ag e m ic he ls en 19 89 b an d th is s tu dy m ic he ls en 19 75 lu nd 1 97 7; d yb kj æ r 19 91 fjerritslev formation n o re co rd s pe rin op ol le ni te s el at oi de s z on e po or re co rd s ca llia la sp or ite s– pe rin op ol le ni te s z on e be rt he ls en 19 79 h oe ls ta d 19 85 r id in g & t ho m as 1 99 2 k op pe lh us & n ie ls en 1 99 4 k op pe lh us & b at te n 19 96 se id en kr an tz e t a l. 1 99 3 an d th is s tu dy ba rr en z on e lll z on e ll z on e l sp he rip ol le ni te s– le pt ol ep id ite s z on e o . a de nt ic ul at a– n . ( n .) sim pl ex z on e ce re br op ol le ni te s m ac ro ve ru co ss us z on e sp he rip ol le ni te s– le pt ol ep id ite s z on e ch as m at os po rit es z on e lu eh nd ea sp in os a z on e lu eh nd ea sp in os a z on e m en di co di ni um re tic ul at um z on e fd pd? pc pb pa fc fb fa n an no ce ra to ps is gr ac ilis z on e n an no ce ra to ps is gr ac ilis z on e d c 4 c 3 h al da ge r sa nd fo rm at io n u pp er pl ie ns ba ch ia n c hr on os tr at ig ra ph y li th ost ra tig ra ph y o st ra co ds sp or es /p ol le n d in of la ge lla te s fo ra m in ife ra pa ly no m or ph s m io sp or es bi os tr at ig ra ph ic al z on at io ns f-iii memberf-iv mb ababcde fi g. 3 . c o rr el at io n o f th e a n h o lt p al yn o m o rp h a n d f o ra m in if er al z o n at io n w ith t h e ch ro n o an d l ith o st ra tig ra p h y an d w ith s el ec te d b io st ra tig ra p h ic al z o n at io n s fr o m p re vi o u s st u d ie s in t h e re gi o n . 590 dev 1961, nannoceratopsis gracilis alberti emend. van helden 1977, and nannoceratopsis senex van helden 1977. this palynomorph assemblage correlates with the perinopollenites elatoides zone from the stenlille boreholes in the danish basin (dybkjær 1991), zone ii from the bagå formation at the hasle klinkerfabrik clay pit, bornholm (hoelstad 1985) and the lower part of the callialasporites–perinopollenites zone from the bagå formation in borehole 107 and the korsodde section, bornholm (koppelhus & nielsen 1994). the zone is also recognised in the rya formation in the øresund area (fig. 3; koppelhus & batten 1996) and from assemblage zone 6 in the neill klinter group, east greenland (koppelhus & dam 2003, this volume). this correlation suggests an aalenian age for zone pc. palynomorph zone pd characterised by perinopollenites elatoides, callialasporites turbatus, c. dampieri (balme) dev 1961, c. microvelatus schulz 1966, c. minus (tralau) guy 1971, densoisporites scanicus tralau 1968, neoraistrickia gristhorphensis (couper) tralau 1964, sestrosporites pseudoalveolatus (couper) dettmann 1963, gleicheniidites conspiciendus (bolchovitina) krutzsch 1959, and g. senonicus ross 1949. this assemblage suggests a correlation with zone d in hobro-1 (bertelsen 1979), with zone iii from the bagå formation (hasle klinkerfabrik clay pit; hoelstad 1985), and with the upper part of the callialasporites–perinopollenites zone, also from the bagå formation on bornholm (bagå beds, section 1; koppelhus & nielsen 1994). the age of the assemblage is considered to be middle jurassic, probably within the bajocian/bathonian interval. palaeomagnetic studies the continuous record of palaeomagnetic anomalies derived from the deep-sea record only reaches back to the late middle jurassic (base of m29 of the top callovian; harland et al. 1990). older reversal chronology therefore relies upon palaeomagnetic polarity studies of individual stratigraphical sections (ogg et al. 1984; steiner et al. 1985, 1987; gradstein et al. 1994). the aim of the present palaeomagnetic study was to contribute to the gradually increasing knowledge of reversals in mesozoic times. this palaeomagnetic study is the first of its kind in the danish area, as previous studies from the area focussing upon pre-quaternary strata are few (e.g. abrahamsen 1994; ali et al. 1994), and none of these were from the jurassic. magnetic methods and results a total of 122 samples were collected for magnetic studies from the cored sections in anholt iii and iv. the samples were obtained by pressing one-inch cylindrical polystyrene holders into argillaceous parts of the stiff sediment. most of the samples are from the toarcian (78 samples), whereas 17 are from the aalenian, and 27 are from the bajocian/bathonian section. all samples were routinely measured in a molspin ltd. spinner magnetometer and subsequently step-wise demagnetised in 5, 10, 20, and 30 mt alternating fields (af) (three orthogonal directions inside metal shields) and remeasured after each step. in order to determine the carriers of the magnetic remanence, eight samples were treated in step-wise increasing pulse-magnetic fields up to 0.6 t (fig. 4). the 159.61 m 189.27 m 249.49 m 252.33 m 121.01 m 160.54 m 121.32 m 186.93 m 0.0 0.1 0.2 0.3 0.4 0.5 0.6 180 160 140 120 100 80 60 40 20 0 field (t) in te ns ity ( m a /m ) fig. 4. irm (isothermal remanent magnetisation) performed on eight samples, which show saturation between 0.2 and 0.3 t, suggesting magnetite to be the carrier of the remanence. 591 irm (induced remanent magnetisation) typically saturates in fields between 0.2 and 0.3 t, indicating that the magnetic carriers are dominated by magnetite rather than hematite or goethite. as the mdf (median destructive field) for most samples is well above 30 mt, this suggests that the magnetic minerals are dominated by singleor pseudosingle-domain magnetites, probably of detrital origin. the nrm (natural remanent magnetisation) intensity is low, typically 1–2 ma/m, with extremes of about 0.4 and 4.5 ma/m (fig. 5). the magnetic declination is not known, as the cores were not oriented in azimuth. the af cleaned inclination (30 mt) is shown in figure 6. the present latitude of anholt is 56.7°n, which is equivalent to a central axial dipole inclination of i0 = 71.8°. however, nearly all the inclinations are much closer to 0° than expected (fig. 6). only 16 out of the 122 samples show inclinations more than ± 30° from 0°. during the jurassic, the danish area had a latitudinal position not significantly south of today (press & siever 1982), indicating that this pronounced inclination shallowing cannot be due to plate drift alone, but rather was caused by either remagnetisation or compaction (abrahamsen 1992). as both polarities are found, secondary magnetisation is unlikely to be the cause. the inclination shallowing in the sediment is thus probably due to compaction. on account of the presence of several major gaps in sampling caused by lithological variation (coarse-grained sections were either not recovered or not sampled for palaeomagnetic studies) as well as due to the dominant shallow inclinations, the reversal chronology based upon the material from anholt is poorly constrained. a tentative interpretation is shown in figure 6, but a more detailed magnetostratigraphical correlation to other palaeomagnetic records such as the detailed sequence from the bajocian–bathonian of south spain (steiner et al. 1987) is not justified at present. 0 1 2 3 4 5 100 120 140 160 180 200 220 240 260 280 300 ba th on ia n– ba jo ci an a al en ia n to ar ci an d ep th ( m b .s .) intensity (ma/m) fig. 5. nrm (natural remanent magnetisation) intensity vs. depth. most nrm intensities are between 0.5 and 2.5 ma/m, with an average of 1.5 ma/m. m b.s., metres below surface. -30-60-90 0 inclination (degrees) 30 60 90 100 120 140 160 180 200 220 240 260 280 300 ba th on ia n– ba jo ci an a al en ia n to ar ci an d ep th ( m b .s .) ? ? ? ? ? fig. 6. suggested polarity interpretation of the jurassic inclination record from the anholt borehole. black, normal polarity; white, reversed polarity; grey, intervals in which polarity is mixed or undetermined; zigzag line, significant intervals in which polarity data is lacking; m b.s., metres below surface. lithology a preliminary sediment description was performed during the drilling operation. a more detailed description was made subsequently, both of cores and cuttings samples, and black-and-white and colour photographs of the cores were taken. x-ray photographs were taken of selected 1 cm thick slices of the core material. immediately after drilling, a sng-log (spectral natural gamma-ray log) was obtained from the anholt iv borehole. preliminary results were published by nielsen (1992). methods the following analyses were carried out: 1. grain-size analysis using wet sieving (> 63 µm fraction) and settling in andreassen tubes (< 63 µm fraction). 2. bulk sediment geochemistry comprising toc (total organic carbon) and sulphur analysis measured in a leco induction furnace, as well as measurements by atomic absorption spectrophotometry (aas) of the content of the main elements: si, al, fe, ca, mg, mn, na, and k. 3. mineralogical composition of both bulk samples and of the clay fraction using x-ray diffraction (xrd). samples for bulk mineralogy were dried in a furnace at 105°c for 16 hours, and ground in a mortar. the powder was mounted in the sample holder without preferred orientation. samples for clay mineralogy were dispersed in distilled water, sieved through a 63 µm mesh, and the < 2 µm fraction was obtained by repeated decantation in distilled water. the < 2 µm fraction was smeared onto glass plates in order to obtain preferred orientation of the clay minerals. the glass plates with the clay film were x-rayed in various conditions: air-dried, ethyleneglycol-treated and after heating to different temperatures (nielsen et al. 1989). 4. some of the samples were heated at a rate of 2°c per minute in a differential thermal analysis/thermal gravimetry (dta/tg) instrument in order to determine the temperatures for oxidation of organic compounds and to release the oh-groups of the phyllosilicate minerals. the change in weight from these processes was determined. the analyses were performed both on bulk samples and on the clay fractions (0–2 µm, 0–1 µm, and 1–2 µm) and were followed by heating to 468°c. this temperature was kept constant for 8 hours and the weight loss from the release of ohgroups as a function of time was recorded. a temperature of 468°c was chosen, because the dta/tg curve indicated that approximately 465°c was the threshold temperature for the dehydroxylation process at the heating rate used. the time in minutes for the release of half of the oh-groups was calculated. 5. the same samples were exposed to heating in steps of 25°c from 450–600°c followed by xrd analysis in order to observe the decomposition of the crystal structure in kaolinite. results and discussion sedimentary facies the clays and sands from 306–201 m are referred to the fjerritslev formation (figs 2, 7; michelsen 1978). in the lower part (c. 306–231 m), the sng-log is rather uniform and the sediment consists of a relatively homogeneous, greyish green clay with a low median grain-size (fig. 8). this interval is equivalent to member f-iii of the fjerritslev formation (michelsen 1978). the lowermost c. 6 m are characterised by a slightly sandy and bioturbated clay (figs 7, 8) with ammonites. pure clay lithologies make up 50–70% of each core, the remainder comprising silty clay layers and fine-grained sand laminae. above this basal unit, the clay is commonly non-bioturbated and contains less than 1% sand, except for very thin laminae of fine-grained sand of inferred storm origin. silty clays form between 1 and 30% of each core, whereas pure clay makes up 70–99% of each core. upwards, especially above 260 m, the storm sand layers become coarser and thicker, up to a few millimetres, and the sand content in the clay layers increases to 4–6% (fig. 8). some sand laminae are cross-laminated. at 231 m, a carbonate-cemented lignite conglomerate (fig. 9) marks the lower boundary of member f-iv of the fjerritslev formation (michelsen 1978). this member continues up to 201 m and is characterised by sandier sediments with more numerous and thicker sand layers intercalated with the clay layers. the sand layers show an upwards increase in the content of very fine pebbles from 1.6–8.5% in the lower part to 11–23% in the upper part (fig. 8). the upper part of the jurassic succession (201–104 m) is referred to the haldager sand formation (michelsen 592 593 100 300 250 200 150 306 fiii m em be r fiv m em be r upper pliens. quat. m id dl e ju ra ss ic lo w er ju ra ss ic to ar ci an h al da ge r sa nd f or m at io n fj er ri ts le v fo rm at io n ba jo ci an /b at ho ni an a al en ia n lithostrat. chronostratigraphy lithology th (ppm) u (ppm) k % total gr log u/th th/k u/k comments 0 0 4 m b.s. 20 5 0 bedding tilted (up to 20°) horizontal bedding calcite-cemented lignite conglomerate siderite concretions siderite concretions siderite concretions pyrite and siderite concretions siderite concretions burrows horizontal bedding concretions silt clay till carbonate lignite clasts gravel coarse-grained sand medium-grained sand fine-grained sand silty sand silty clay fig. 7. integrated log combining lithostratigraphical and lithological data with the spectral natural gamma-ray log (sng). quat., quaternary; pliens., pliensbachian; m b.s., metres below surface. 594 1978). this part has a markedly higher content of sand with brown coal fragments. the lithology of the haldager sand formation shows a cyclic variation which is also evident on the gamma-ray log (fig. 7). intervals (1–10 m thick) of medium-grained sand with lignite fragments alternate with 1–5 m thick units of silty clay with high toc contents (fig. 2). lignite layers, a few centimetres thick, also occur locally. the observed variations in lithology might be explained by lateral migration of channels in a delta-plain or floodplain setting. sng-log analysis the thorium (th), uranium (u), potassium (k), and integrated gamma-ray signals from the spectral natural gamma-ray log (sng-log; fig. 7) obtained from the anholt iv borehole show some significant variations throughout the jurassic section. the main carrier of potassium is believed to be phyllosilicates (clay minerals and mica), as almost no potassium feldspar was observed. uranium is normally concentrated in layers rich in organic matter but may also be present in some heavy minerals. the th-content is somewhat greater in kaolinite than in other clay minerals, but some heavy minerals contain relatively high concentrations of th (see below). the log signal is very uniform in member f-iii of the fjerritslev formation. the only exceptions are characteristic negative peaks where carbonate and pyrite concretions and sand laminae are present (fig. 7). the lower gamma-ray values probably indicate a lower concentration of th, u and k in the carbonate concretions than in the surrounding compacted clay, suggesting that carbonate cementation took place shortly after deposition and prior to significant concentration of k, th and u due to compaction of the clays. the log signature of member f-iv is markedly different from that of the haldager sand formation, mainly due to variations in lithology. the fine-grained layers are enriched in k-containing phyllosilicates and u-containing organic matter relative to the coarser layers. part of the th content is probably also related to kaolinite particles, which are more common in the finer size fractions. uranium and thorium anomalies in the sng-log some thin layers in member f-iv of the fjerritslev formation and in the lowermost part of the haldager sand formation (around 195 m) have higher concentrations of th, and to some extent also u (fig. 7). the smaller th and u anomalies are all located within the coarsening-upwards marine member f-iv, whereas the two strong anomalies are located within the lower levels of the generally non-marine haldager sand formation (environmental interpretations from michelsen 1978, 1989a). the k content is constant or even relatively low in these layers (see also the element ratios in fig. 7). 10 000 1000 100 10 1 90 115 140 165 190 215 240 265 290 315 m ed iu m g ra in s iz e (µ m ) depth (m b.s.) fjerritslev fmhaldager sand fmquat. fig. 8. median grain-size values; note the logarithmic scale on the y-axis. some reservations should be made concerning the results of the grain-size analysis of the f-iv member (231–201 m), as no cores were recovered from this interval. the sediments are only slightly consolidated, and the more fine-grained parts might therefore have been incorporated in the drilling fluid. m b.s., below surface. the base-level of the gamma-ray signal is relatively high in the clay lithologies of the fjerritslev formation, but considerably lower in the haldager sand formation. low gamma signals are normal for sand layers. ditch cuttings samples and samples from the very few cores from these low gamma-ray intervals are dominated by sand thus supporting the log interpretation. however, the very poor core recovery in this section of the borehole did not allow very precise sampling of undisturbed th-rich intervals, so the reason for the high concentrations of this element is unclear. similar peaks of th and u are found in jurassic, cretaceous and miocene deposits in other danish boreholes (korsbech & nielsen 1991; korsbech 1992). the same phenomenon is seen in sections with excellent core recovery of lower cretaceous fine-grained sand in the skagen-3 borehole. here, the coarse silt fraction contains a considerable proportion of heavy minerals (mainly zircons; h. friis, personal communication 1992). the same observation has been made by conner & kelland (1975a, b) in a jurassic sandstone in a north sea well; zircons and other heavy minerals such as monazite, rutile and epidote commonly contain a substantial amount of th, and it was suggested that this might be the source for the relatively high content of this element. the same explanation may be applicable to the jurassic in the anholt cores. an examination of some older borehole logs indicates that layers enriched in heavy minerals may be a common phenomenon in jurassic sandstones in denmark. although sng-logs were not run in these old boreholes, we propose that some of the peaks observed on natural gamma-logs in the jurassic and lower cretaceous sediments of the frederikshavn city-1 and børglum-1 boreholes are due to a high content of heavy minerals. larsen (1966) detected several samples with zircon contents of up to 60% of the non-opaque, nonmicaceous heavy fraction. higher gamma signals are normally found in these layers. the concentrations of th and u between 198 and 190 m are shown on figure 10; note that the u and th peaks are displaced slightly (5–10 cm) relative to one another. the scales were chosen in order to equate the heights of the th and u peaks at 193.9 m. the ratio between the th and u concentrations here is 5:1. in contrast, the th:u ratio for the peaks at 195 m is about 10:1; note that minor fluctuations occur within these peaks at 195 m. this variation in the th:u ratio may imply that some separation of the heavy minerals has taken place, probably based on differences in grain size and density. bulk mineralogy the mineralogical composition of the jurassic succession of the anholt borehole is relatively simple; it comprises only stable, detrital minerals and a few authigenic/diagenetic minerals. the detrital minerals are quartz, clay minerals (including mica) and very little feldspar; the authigenic/diagenetic minerals are pyrite, calcite and siderite. authigenic feldspars have been 595 fig. 9. core photograph of part of the interval from 232.6–231.1 m b.s., fjerritslev formation, f-iii member, uppermost toarcian. opposing faces of a single slabbed core piece are illustrated. the dark clasts are lignite fragments, up to 7 cm in length, in a matrix of calcite-cemented clayey, silty sandstone. width of core is 10 cm. 596 reported in middle jurassic sediments elsewhere in the north sea basin (nielsen & friis 1985), but were not observed in this borehole. the mineralogical variation mainly reflects the grainsize variation and the location of pyrite and carbonate concretions of predominantly sideritic composition. the boundary to the overlying quaternary sediments is clearly marked by a distinct increase in feldspars and amphiboles, i.e. minerals with a relatively low resistance to chemical weathering. clay mineralogy the non-clay minerals of the < 2 µm fraction are dominated by quartz, but the quartz content is much lower in fractions below 1 µm. chlorite is only present in small quantities with a slight tendency to decrease upwards (fig. 11), presumably due to the diagenetic origin of this mineral. illite is a common mineral although the measured illite content in the coarser layers is presumed to also include muscovite (identified macroscopically), as it is difficult to distinguish between these two very similar minerals by xrd. kaolinite is the dominant clay mineral in the coarser sediments. in sands from the haldager sand formation, macroscopic examination revealed a white powder, identified as pure kaolinite by xrd. in finer grades of the clay fraction, the kaolinite content is notably lower, due to its normally larger grain-size relative to other clay minerals. smectites are present in the lower, fine-grained layers (figs 11, 12), but decrease in proportion up-section, rel0 0 190 198 50th (ppm) 10u (ppm) d ep th ( m b .s .) th u fig. 10. thorium (th) and uranium (u) concentrations between 198 and 190 m b.s. the relative concentrations (th/u) differ for the two peaks, probably reflecting sorting processes during sedimentation. m b.s., below surface. 100 90 80 70 60 50 40 30 20 10 0 95 104 201 305 % ( cu m ul at iv e) quat. haldager sand fm fjerritslev fm kaolinite chlorite illite smectite fig. 11. clay mineralogy of the < 2 µm fraction. sum of clay minerals = 100%. the analyses are displayed in stratigraphic order but the horizontal scale is not directly proportional to depth; the depth to the main lithostratigraphic divisions is indicated (in metres below surface). 597 ative to the total clay fraction, corresponding to the overall increase in grain size up-section (fig. 11). the presence of smectite in the fjerritslev formation should be noted as smectite is absent in jurassic and lower cretaceous sediments in most wells in the danish basin and the danish central graben (nielsen 1979, 1985; clausen 1982; nielsen & friis 1985; schmidt 1985a; krabbe 1986; lindgreen 1991). we believe that the dominance of illite and the presence of illite/smectite mixedlayer minerals in jurassic sediments in most boreholes, excepting anholt, is the result of substantial diagenetic illitisation of original smectites due to greater burial depths and thus exposure to higher temperatures. the transformation of smectites to illites via mixedlayer phases is primarily dependent on burial history, temperature and time (dunoyer de segonzac 1970). the transformation is thus seen at different depths and ages in different basins. the process mainly happens within a temperature range of 80–115°c, normally corresponding to a burial depth of 2–4 km in the north sea (środoń & eberl 1984). values of these magnitudes have never been reached for the anholt sediments. the 4. 76 50 5. 00 87 5. 29 86 5. 61 14 5. 75 71 7. 15 54 10 .0 77 7 14 .0 04 6 16 .7 46 3 18 .6 32 7 20 .6 92 5 22 .6 58 3 33 .9 07 6 5. 02 90 7. 24 62 9. 63 58 10 .1 59 2 13 .2 70 3 14 .6 67 4 16 .8 91 2 18 .3 54 4 22 .1 24 6 30 .4 39 6 32 .6 14 4 ill ill chl chl sm ka/chl heat-treated eth.-treated fig. 12. x-ray diffractogram of the clay fraction of a sample from the fjerritslev formation (294.24–294.21 m b.s.). heattreated, x-rayed after heating to 500°c for 1 hour; eth.-treated, x-rayed after treatment with ethyleneglycol vapour for 16 hours at 60°c. reflection peaks: chl, chlorite; ill, illite; ka, kaolinite; sm, smectite. the figures represent crystal lattice spacings in ångstrøms. 598 illitisation process starts in sediments with vitrinite reflectance values of approximately 0.5 %ro, possibly already at 0.40–0.45 %ro (pevear et al. 1980), and burtner & warner (1986) proposed on the basis of rock-eval analysis that the process takes place in sediments with tmax values between 430° and 440°c, possibly already at 425°c. in deeply-buried jurassic successions, such as in the børglum-1 and uglev-1 wells, the vitrinite reflectance values for sediments of the fjerritslev formation are 0.40–0.50 %ro (schmidt 1985b), i.e. considerably higher than the values from anholt (0.25–0.40 %ro, see below). the percentage of smectite normally increases with decreasing particle size and towards the basin centre. because of the marginal location of the anholt borehole, it is possible that the sediments contain less detrital smectite than contemporaneous sediments in other wells in the danish basin prior to diagenetic alteration. geochemistry the toc values are shown in figure 2 and the s, fe2o3 and k2o values in figure 13. peak values of toc in the haldager sand formation are related to brown coal layers. intermediate values of toc are usually found in brownish/black layers dominated by silt-sized particles, normally with relatively abundant muscovite, or in sand layers with scattered lignite clasts, such as in the middle of the fjerritslev formation section. low toc values are mainly restricted to layers dominated by clay or coarse-grained quartz sand. the k2o content generally reflects the amount of illite, whereas s and fe2o3 reflect the pyrite and to some extent the siderite contents. the percentage of fe, excluding that incorporated in pyrite or siderite, is relatively uniform. the analyses were used to calculate a normative mineralogical composition, which was then compared with the xrd mineralogy and used to improve the quantification of the bulk mineralogy from the xrd data. the anholt borehole is well-suited to such a procedure as: (1) many core-samples have been analysed, (2) the number of different minerals is limited, and (3) the minerals present are rather stable. differential thermal analysis (dta) and step-heated xrd the result of the step-wise heating of the sediment for one hour followed by xrd is shown in figure 14. in samples from the top of the jurassic section down to a depth of about 230 m, the 7å kaolinite peak on x-ray diffractograms is totally destroyed by heating to 500°c for one hour; indeed, kaolinite from sand layers in the haldager sand formation is already destroyed at a temperature of 450°c. from 230 m down to the base of the borehole at 306 m, the kaolinite structure is only partly destroyed by heating for one hour. the higher the heating temperature and the longer the sample is exposed to this temperature, the greater the destruction. measurements of different size fractions clearly demonstrated that larger grains were more heat resistant than smaller grains, probably due to a greater proportion of grains with crystal lattice imperfections amongst the smaller size fractions; such size-dependent thermal stability has been noted in many previous studies (e.g. norton 1939). with the exception of a single sample at about 250 m, there is a slight tendency for kaolinites in member f-iii to be more heat resistant with increasing depth, even if the total amount of kaolinite decreases (figs 11, 15) and the particle size of the mineral grains decreases (fig. 8). figure 15 shows the relationship between core depth and time to destruction of half of the kaolinite particles at 468°c, i.e. the release of oh-groups corresponding to half of the weight loss observed at this temperature. generally, the time needed for this destruction increases with increasing depth. this tendency towards a down-section increase in thermal resistance of kaolinite has been observed elsewhere in mesozoic sediments of the north sea basin. holdridge & vaughan (1957) found the temperature for the onset of dehydroxylation to be 477°c for kaolinites (average of 48 samples) and 522.5°c for pure dickite (average of 18 samples) through measurement on differential temperature analysis (dta) equipment at a heating rate of 10°c per minute. the lower threshold temperature observed in these anholt data (468°c), is probably due to the lower heating rate (2°c/min.) used in the present study. the crystal order presumably increases with the influence of progressive diagenesis. in deeper diagenesis/incipient metamorphism, kaolinites may convert to dickite and further to nacrite with the same chemical composition as kaolinite, but with a higher degree of order in the crystal lattice (dunoyer de segonzac 1970). the results of the analysis of the thermal stability of kaolinites from the anholt borehole indicate that the thermal stabilisation of kaolinites during progressive diagenesis is a gradual process. the process seems to initiate at a relatively early diagenetic stage; thermal stabilisation is detectable at diagenetic stages reached 599 at a depth of approximately 1 km. it also seems possible to differentiate the thermal stability of kaolinites within a depth range of 100–200 m. kaolinite in sands of the haldager sand formation is commonly seen as a whitish powder on the quartz grains. the very low thermal resistance of this kaolinite might be explained by a later, probably post-uplift, formation in the pore spaces of the sand. the formation of kaolinite is possibly due to reactions between a migrating diluted acidic porewater and feldspar grains (nielsen & friis 1985). organic petrology and geochemistry a total of 46 samples from the anholt borehole were selected for toc and rock-eval screening. the samples were selected mainly from the darker grey, brownish or black sediment levels where higher organic contents were expected. from the screening results, 23 samples were selected for organic petrographical analyses, i.e. measurements of vitrinite reflectance and fluorescence, and compositional description. 35 30 25 20 15 10 5 0 90 110 130 150 170 190 210 230 250 270 290 310 depth (m b.s.) pe rc en t quat. haldager sand fm fjerritslev fm k2o fe2o3 s fig. 13. the content (%) of k2o, fe2o3, and s determined from the bulk geochemical analysis. quat., quaternary; m b.s., metres below surface. fig. 14. peak ratios (peak height of heated samples/peak height of ethyleneglycol-treated samples) for 7å kaolinite heated to 500°c, 525°c and 550°c. m b.s., metres below surface. 220 0.00 0.02 0.04 0.06 0.08 0.10 r at io 0.12 0.14 0.16 230 240 250 260 270 depth (m b.s.) 280 290 300 310 500°c/glycol 525°c/glycol 550°c/glycol 600 toc and rock-eval screening most of the samples are very poor in total organic carbon (toc). the hydrocarbon potential (hc) is also low (s2 values are low; table 1). s2 values (s2: hydrocarbons generated during rock-eval pyrolysis) of < 0.5 mg hydrocarbon (hc)/g rock are considered to give unreliable tmax values (tmax: temperature of maximum hydrocarbon generation during rock-eval pyrolysis). the highest toc values are from the upper sandy section of the haldager sand formation from a depth of about 190–105 m, whereas the values from the underlying greenish clays of the fjerritslev formation are low to very low. the varying content of organic/coaly debris in the sediments is confirmed by the wide range of variation in toc values. the highest value (38%) was obtained from a thin lignite layer. the tmax varies in the range 410–433°c, with most of the values around 420–425°c (table 1). the lowest value of 387°c from the brown coal bed is considered to be unreliable. the highest value of 433°c is most likely due to the presence of slightly oxidised or reworked material and not due to maturation. the dominant tmax values of about 420–425°c are within the expected range for sediments with the measured vitrinite reflectance values (see below) of about 0.4 %ro (tissot et al. 1987). the hi (hydrogen index) below or just above 100 mg hc/gtoc (table 1) is typical of kerogen types iii–iv. only one sample yielded a higher hi value (276 mg hc/g toc). coal petrography the samples selected for coal petrographic description were embedded in cold synthetic resin, and were ground and polished after hardening. all samples were studied in normal reflected light combined with blue lightinduced fluorescence. the organic material was described and the maceral composition was estimated semi-quantitatively in the following groups: vitrinite, pseudovitrinite, inertinite and liptinite. the main focus was on the degree of oxidation and the liptinite composition. point counting was not performed; the results of the coal petrographic analysis are given in table 2. the organic material in all samples is, in general, dominated by terrestrial material. most samples are characterised by humic material (vitrinite) showing varying degrees of oxidation and/or reworking (pseudovitrinite and inertinite). samples 3 and 8, from 121.6 m and 170.9 m respectively, are composed entirely of large fragments of ulminitic vitrinite with almost no indication of oxidation. in contrast, sample 10 (184.48 m) contains humic material composed almost entirely of oxidised inertinite, mainly tiny small fusinite fragments, or pseudovitrinite. in general, the liptinitic content is low, but the identifiable part shows a high abundance of macerals of terrestrial origin, such as sporinite, resinite and cutinite, sometimes even with possible fluorinite. only very few algae were identified. in the upper part of the section, a few of the botryococcus-type occur, indicating deposition in (or derivation from) fresh to brackish water, and a tasmanites was observed in sample 10 (184.48 m). marine organic material, such as tiny fragments of dinocysts or acritarchs, is present only in the lowermost 100 120 140 160 180 depth (m b.s.) m in ut es a t 46 8° c k ao lin ite ( % ) 200 220 240 260 280 60 50 40 30 20 10 0 1/2 weight at 468°c kaolinite 100 90 80 70 60 50 40 30 20 10 0 fig. 15. the proportion of kaolinite (%) relative to the total clay mineral content of the clay fraction (< 2 µm) and the time taken in minutes for the release of half of the oh-groups in kaolinite at 468°c, measured by differential thermal gravimetry. m b.s., metres below surface. part of the succession. a very small amount of bituminite was also recognised in these samples. sample 22 (300.88 m) has the highest proportion of what may be considered to be marine material. the organic material in the uppermost part of the fjerritslev formation is of very poor quality compared to data from other wells in the danish area (thomsen et al. 1987). vitrinite reflectance and fluorescence data reflectance measurements were performed on the same polished specimens as used for description of the organic composition. the photomultiplier attached to the ‘zeiss photomicroscope’ was calibrated by measurements on a polished optical glass standard of constant known 601 1 113.6 0.30 52 2 119.82–119.85 11.00 417 1.08 12.17 0.08 1.10 111 0.33 51 3 121.59–121.60 38.00 387 5.70 31.07 0.16 3.05 82 0.32 63 4 124.6–127.6 3.74 427 0.19 3.15 0.06 0.28 84 0.33 39 130.6–133.6 2.02 426 0.11 1.25 0.08 0.11 62 5 133.6–136.6 2.11 423 0.07 0.97 0.07 0.09 46 0.37 49 137.4–139.6 0.17 424 0.01 0.01 0.50 0.00 6 142.6–145.6 0.17 404 0.03 0.05 0.38 0.01 30 145.6–147.1 1.35 425 0.01 0.11 0.08 0.01 8 147.1–148.6 2.43 424 0.05 0.91 0.05 0.08 37 6 149.90 0.70 426 0.12 1.94 0.06 0.17 276 0.35 50 154.6–157.6 2.44 423 0.03 0.23 0.12 0.02 9 7 159.59–159.74 2.40 424 0.13 1.74 0.07 0.16 73 0.35 52 8 170.8–170.90 5.89 410 0.35 2.67 0.12 0.25 45 0.32 55 9 175.55 1.81 423 0.03 0.79 0.04 0.07 44 0.43 43 184.09–184.10 0.77 427 0.01 0.24 0.04 0.02 31 10 184.80 9.50 427 0.46 5.94 0.07 0.53 63 0.34 7 184.98 1.73 421 0.07 1.32 0.05 0.12 76 11 187.82 1.50 418 0.03 0.59 0.05 0.05 39 0.34 30 12 189.65 0.49 431 0.00 0.07 14 0.32 27 194.6–197.6 0.12 0.01 0.00 13 206.6–209.6 0.63 424 0.03 0.36 0.08 0.03 57 0.34 30 209.6–210.3 0.38 425 0.03 0.15 0.17 0.01 40 212.6–215.6 0.31 422 0.03 0.13 0.19 0.01 42 14 218.6–221.6 0.31 425 0.05 0.13 0.28 0.01 42 0.32 50 221.6–224.6 0.28 422 0.05 0.11 0.31 0.01 40 15 230.53 1.19 422 0.03 0.50 0.05 0.05 50 0.33 51 234.36 0.16 431 0.00 0.07 44 16 239.28 0.72 429 0.01 0.42 0.02 0.04 58 0.31 14 244.15 1.20 427 0.01 1.31 0.01 0.11 109 17 249.68 0.24 429 0.00 0.17 69 0.29 8 255.16 0.22 432 0.00 0.07 32 260.59 0.21 434 0.00 0.13 61 264.20 0.28 420 0.00 0.17 60 18 270.70 0.50 420 0.01 0.25 0.04 0.02 50 274.80 0.19 419 0.00 0.07 38 19 280.09 0.17 414 0.01 0.07 0.13 0.01 41 285.70 0.20 416 0.00 0.07 36 20 290.08 0.45 420 0.01 0.30 0.03 0.03 67 0.31 8 21 295.08 0.42 416 0.01 0.25 0.04 0.02 59 22 300.08 1.08 420 0.05 1.01 0.05 0.09 94 0.34 38 23 304.08 0.66 423 0.01 0.34 0.03 0.03 52 0.35 36 sample depth (m) toc (wt%) tmax (°c) s1 (mg hc/g rock) s2 (mg hc/g rock) pi pc (wt%) hi %ro n table 1. rock-eval screening results and vitrinite reflectance toc : total organic carbon. tmax: temperature of maximum hydrocarbon generation during pyrolysis. s1: thermally-extracted hydrocarbons. s2: hydrocarbons generated during rock-eval pyrolysis. pl: production index [s1/(s1+s2)], derived from rock-eval pyrolysis. pc: pyrolysable carbon [0.083 (s1+s2)]. hl: hydrogen index [100 x s2/toc]. %ro: vitrinite reflectance. n: number of measurements. reflectance. the reflectance measurements are reported as the percentage (%ro) of the incident light reflected from the plane surface of the vitrinite surface when immersed in oil. the abundance of vitrinite particles is good for reflectance measurements, even in the lean samples, as most samples are dominated by humic material. in general, as many measurements as allowed by the available material were performed, and the quality of the data was noted. the average values shown in table 1 are based on reflectance measurements that were considered to be reliable. the most reliable data are considered to be from the coaly samples where the particles are so large that a detailed maceral identification is possible, and therefore the results are more reliable compared with samples containing fine-grained dispersed organic matter. in sample 3 (121.6 m), the organic material is represented entirely by large particles of woody tissue and the measurements were performed on texto-ulminite. the range of values is very narrow (0.25–0.40 %ro) with an average of 0.32 %ro from a total of 62 measurements. reflectance values from corpohuminite, which is also common, have the same range and almost the same average (0.31 %ro). the range of values in the other samples is somewhat larger, though still between 0.2 and 0.5 %ro. the best data quality is generally obtained from the upper coaly samples, whereas the quality decreases in some of the lowermost samples where the particle size of the 602 table 2. composition of the organic material v itr in ite ps eu do vi tr in ite in er tin ite li pt in ite sample maximum depth (m) organic content maceral composition liptinite composition d in of la ge lla te s/ ac ri ta rc hs ta sm an ite s bo tr yo co cc us bi tu m in ite li pt od et ri ni te sp or in ite c ut in ite fl uo ri ni te r es in ite 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 113.6 119.85 121.60 127.60 136.60 149.90 159.70 170.90 175.55 184.48 187.82 189.65 209.60 221.60 230.53 239.28 249.68 270.70 280.09 290.08 295.08 300.88 304.08 moderate high very high moderate moderate moderate moderate high moderate high moderate low low low low low low very low very low low low low low xxxx xxx xxxxx xx x xxx xxx xxxxx xxx x xxxx xxx xxx xxxx xx xxxx xxx xxxx xxx x x xx xxx x xx xx xxx xx xx xx xx x x xx x xx x x x x xx xx x xx x x xx xx x x x x xxxx x x x x xx x x x xx xx x x x x x x x x x xx x xx xxx x x x x x x x x x x x x x xx x x x x xx x x x xx xx x xx x x x x x x x x x x xx xx x x xx xx x xx x x x x x xx x: abundance, ranked in terms of relative abundance (x–xxxxx). : observed. vitrinite is smaller and the morphological affinity of the vitrinite less obvious. the poorer quality is indicated by the wider range of the reflectance data. in samples 18, 19 and 21, it was not possible to perform measurements but the range of reflectance data in table 1 is considered reliable, due to the relatively large number of measurements on each sample. spectral fluorescence measurements were performed on a leitz orthoplan microscope equipped with a hamamatsu photomultiplier and a ploemopak for measurements and observations. the spectra were measured and their q value recorded (q = relative fluorescence intensity at 650 nm / relative fluorescence intensity at 500 nm). it was only possible to make fluorescence measurements on a few samples containing alginite, i.e. samples 4, 6, 7, 10, 21 and 22. the mainly yellowish-green fluorescence colours of the algal bodies indicate their degree of immaturity and correspond to their vitrinite reflectance values. all average q values are between 0.4 and 0.5, corresponding to the measured vitrinite reflectance values (schmidt 1989). uplift the increase in vitrinite reflectance down-section is the result of chemical processes controlled by time and temperature; the increase is linear with respect to time and exponential with respect to temperature (waples 1985). the changes to the chemical structure of the vitrinite are irreversible, and therefore the vitrinite reflectance reflects the maximum temperature to which the sediment has been exposed during burial. uplift will consequently not cause any change in the vitrinite reflectance. a difference between measured reflectance values from a certain depth and estimated reflectance values for that depth can therefore be used as an indication of the magnitude of uplift. a reflectance value of 0.20 %ro is normal for zero coalification at the depositional surface (teichmüller & teichmüller 1979). reflectance values above 0.3 %ro may first be expected at a depth of at least 1000 m in areas of the north sea with no uplift but a similar temperature history (hansen & thomsen 1990; schmidt 1990). the range of reflectance values slightly above 0.3 %ro suggests that some uplift and erosion has taken place. the uplift is estimated to have been about 1000 m; a comparable value (975 m) was obtained by petersen et al. (2003, this volume) who utilised the reflectance data reported here but applied a coalification curve derived from the danish basin. study of the organic material has thus indicated that: 1. the degree of diagenetic maturation is relatively low, but higher than expected for sediments at a depth range of approximately 100–300 m. 2. the ranges of tmax, vitrinite reflectance and q values suggest that the deposits have been exposed to temperatures that are commonly present at c. 1000 m depth. 3. the hi values and the coal petrographic interpretation indicate a terrestrial origin for most of the organic material. 4. the content of marine organic material is very low, and restricted to the lowermost few metres of the cored section. discussion palaeoenvironment nodosarid foraminifera, such as those found in the lowermost part of the borehole in foraminiferal zone fa from 306–304 m, referred to member f-iii of the fjerritslev formation, were typical of normal marine shelf conditions during the jurassic period (e.g. nagy et al. 1990). the nodosarid assemblage combined with the abundant dinoflagellate cysts of low diversity, the occurrence of ammonites, fish remains, bioturbation, the alternating clay and fine-grained sand lithology, and the low toc content (below 1%; fig. 2) suggest that a marine, inner shelf environment with well-oxygenated bottom conditions prevailed during the late pliensbachian (seidenkrantz et al. 1993; nagy & seidenkrantz in press). storm sands are common and relatively coarse-grained in the lowermost levels of the borehole (306–300 m). in contrast to this basal interval, the storm sands become less abundant and finer-grained above 300 m, although their abundance and grain-size varies, with a slight tendency to become more abundant and coarser upwards between 300 m and 260 m. in the uppermost pliensbachian – toarcian, bioturbation is less prominent and the clay-rich sediment contains an upwards increasing proportion of storm sand laminae, especially from 260 m. the thickness and grain size of the sand laminae also increase upwards, suggesting a gradual decrease in water depth through the toarcian. this is supported by the low abundance of 603 marine dinoflagellates, decreasing up-section, and the low-diversity agglutinated foraminiferal fauna (seidenkrantz et al. 1993; nagy & seidenkrantz in press). such foraminiferal assemblages may reflect low salinity, oxygen deficiency, turbulent conditions or rapid sedimentation and were characteristic of deltaic or prodeltaic environments in the jurassic (løfaldli & nagy 1980; nagy et al. 1988, 1990, 1995). thus, the dominance of ammobaculites, bulbobaculites, kutsevella and haplophragmoides, as well as the low diversity (cf. nagy et al. 1984, 1990; nagy 1992) suggest that the uppermost pliensbachian – toarcian sediments at anholt record progressive shallowing of the environment. the presence of a few ammonite fragments and some dinoflagellate cysts show that marine influence persisted. the absence of ostracods, the scarcity of burrows and the relatively common occurrence of haplophragmoides may reflect reduced oxygen conditions at the sea floor. during the following main part of the toarcian succession (upper palynomorph zone pb, 288–231 m) the sparse agglutinated foraminifera (ammobaculites and bulbobaculites) and the marine dinoflagellate cysts gradually disappear. upwards, storm-generated sand layers become more frequent, thicker and coarser and, especially above 260 m, these sands have erosional bases and show cross-lamination. this supports the interpretation of a progressive shallowing, indicated by the palaeobiological data. in the upper part of the interval, the presence of scattered lignite clasts and the increased proportion of reworked palynomorphs suggest an increasing supply of particles from a shallower part of the coastal zone, temporarily exposed to erosion, and/or from a neighbouring coastal plain. the coal petrographic interpretation indicates that most of the organic material has a terrestrial origin. member f-iv of the fjerritslev formation (231–201 m) is broadly referred to the aalenian stage. there are only few cores in this section, but the presence of layers with relatively low total gamma-ray signals, indicates that substantial parts of the unit consist of sand and silty sand and were deposited under higher energy conditions than the sediments below. the palynomorph assemblage contains a significant proportion of freshwater algae (palynomorph zone pc), which, combined with the occurrence of brackish-marine dinoflagellates and acritarchs, indicates a brackish-marine influence in the deltaic environment. the remaining part of the jurassic section, the uppermost aalenian – bajocian/bathonian haldager sand formation (201–104 m), represents a delta plain environment with only occasional brackish water influence in the lower part. this is shown by the absence of foraminifera and dinoflagellates, by the occurrence of acritarchs only in the lower part, and by the presence of brackish to freshwater algae throughout the section. the coal petrographic analysis indicates a marked dominance of terrestrial material, commonly with a varying degree of oxidation of the organic material. the cyclic alteration of sand, pebbly sand, clay, and lignite supports the delta plain interpretation. in the lowermost haldager sand formation, however, discrete thin finegrained sand beds are significantly enriched in th and u (fig. 7), probably reflecting concentration of heavy minerals under high-energy (beach?) conditions. comparison with the central danish basin our data indicate that at anholt most of the toarcian period was characterised by decreasing water depths. in the lower part of the toarcian section, evidence of shallowing is less pronounced, but is very distinct in the mid-toarcian section. in the central danish basin, much of the toarcian was characterised by an increase in water depth (michelsen 1978, 1989a), which may be correlated with a global sea-level rise (hallam 1988). the sea-level rise led to the establishment of stagnant bottom conditions (pedersen 1986; michelsen 1989a). in the central part of the danish basin, the shallowing of the basin first began in the late toarcian (michelsen 1989a, b). although the biostratigraphic data from anholt does not prove the presence of a complete toarcian section, we believe that most of the toarcian stage is represented, at least within the section of complete recovery below 248 m, as no indications of hiatuses have been observed here. it is deemed likely, therefore, that the regressive, shallowing event was initiated earlier at anholt than in the central parts of the danish basin. we suggest that such an early regressive tendency at anholt may be explained by the palaeogeographical position of the site. anholt is located in a marginal position relative to the danish basin and it is possible that progradation of the delta plain that characterised the entire danish basin in the middle jurassic (michelsen 1978), may have influenced anholt earlier than elsewhere in the basin. another possible explanation may be sought in local tectonic uplift in the fennoscandian border zone, but we have no other indications of such an early jurassic uplift in the area. the boundary between the fjerritslev and haldager sand formations in the danish basin has previously been considered to coincide with the lower–middle jurassic 604 boundary (michelsen 1978, 1989b) although michelsen & nielsen (1991) suggested that the uppermost part of the fjerritslev formation may be aalenian in age. our analyses have shown that member f-iv of the fjerritslev formation was deposited in aalenian time. at anholt, the toarcian–aalenian boundary (i.e. the lower–middle jurassic boundary) is thought to correspond broadly to the boundary between members f-iii and f-iv of the fjerritslev formation (fig. 3). post-depositional history and diagenesis the depth interval with vitrinite reflectance data is only about 200 m thick and a depth-related trend in reflectance is not evident. nevertheless, a tentative estimate of uplift has been derived, suggesting that about 1000 m of overburden are missing; this figure is compatible with that derived by petersen et al. (2003, this volume). this indicates that the maximum post-depositional subsidence of the jurassic sediments at anholt was about 1100–1300 m. this value is not very accurate, but the range seems to fit well with data obtained from other methods for quantification of uplift and erosion such as shale compaction, where sonic and density logs are used to estimate the magnitude of uplift (strømnes 1991; japsen 1992; jensen & michelsen 1992; jensen & schmidt 1992, 1993; lykke-andersen et al. 1993). the range of uplift values is compatible with values estimated from the northernmost part of jylland (japsen 1992; jensen & michelsen 1992; jensen & schmidt 1992, 1993). together they define a regional uplift trend almost parallel to the fennoscandian border zone. the regional trend in uplift for the danish area shows that no uplift or erosion has occurred in the central graben, whereas increasing uplift took place in a northeasterly direction towards the norwegian–swedish coast, reaching a maximum of about 1500 m. figure 16 shows how the data from anholt fit into this overall trend. though the magnitude of uplift estimated from differ605 0 m 500 m 1000 m 800 m 1200 m 1500 m ? ? ? ? neogene depocentre 1.0 1.5 2.0 n dk g nl 58°n 4°e 6°e 8°e 10°e 12°e 57°n 56°n anholt denmark sweden norway iso-uplift contours of the regional neogene uplift neogene depocentre: two-way time (sec.) to base of neogene major fault (reverse, normal) well used in this study, magnitude of uplift in metres iku shallow borehole iso-uplift contours of the late cretaceous – early tertiary inversion 0 200 400 1000600 800 ■■ ■■ ■■ ■■ ■ ■ ■■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ▲ ▲ ■■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ▲ ▲ ■ ■ ■■ ■ ■ ■■ ■■ ■ ■ 400 900 700 700 800 800 800 800 1000 1200 1200 1200 1700 2000 1000 1000 1500 1500 100 km ▲ ■■ 1.0 fig. 16. iso-uplift map of the danish basin. modified from jensen & michelsen (1992). north sea sectors: dk, denmark; g, germany; n, norway; nl, the netherlands. 606 ent methods differs somewhat (japsen 1992, 1993), the overall trend in regional uplift superimposed by inversion along the sorgenfrei–tornquist zone is clear. anholt is situated close to the sorgenfrei–tornquist zone which was inverted during late cretaceous – early tertiary times. this zone is characterised by very high total uplift values reaching about 2000 m (e.g. for the terne-1 well; b.j. schmidt and l.n. jensen, personal communications 1994) representing an addition of 600–700 m to the regional uplift. however, the uplift at anholt does not reach a magnitude which demands a contribution from late cretaceous – early paleocene uplift within the inversion zone (fig. 16). a structural map of the top chalk group (ter-borch 1987) shows that upper cretaceous sediments are found at depths of 150–250 m below sea level both to the ne and sw of anholt. on anholt, lower–middle jurassic sediments occur at this depth, indicating differential uplift of the anholt site of at least this magnitude relative to the areas to the ne and sw. this uplift might have taken place during the inversion of the sorgenfrei–tornquist zone. thus, part of the uplift probably took place during the late cretaceous – paleocene inversion, but the main uplift of the area took place in oligocene–recent times as part of the regional uplift of scandinavia. the mineralogical composition, dominated by detrital quartz, muscovite/illite and kaolinite, with a subordinate smectite contribution increasing down-section, indicates that the source areas for the jurassic sediments had been exposed to intense weathering. diagenetic minerals such as chlorite are present in the lower part, but the presence of smectite demonstrates that the illitisation process, seen in most other wells in the danish basin, has not been significant. this is in accordance with the relatively minor post-jurassic subsidence of the anholt area. the thermal stability of the kaolinite increases down-section. the well-crystallised authigenic kaolinite in the sand layers of the haldager sand formation is less thermally stable than the detrital kaolinite. this indicates that the authigenic kaolinite formed after or during the late cretaceous – tertiary uplift and thus was not thermally stabilised during the subsidence-related progressive diagenesis. the spectral natural gamma-ray log has enabled better definition of the lithological boundaries, especially in sections with poor core recovery. in the transition between the marine clays and the non-marine sands, very high concentrations of th and u are present in relatively thin and well-defined layers. the origin of these layers is not yet fully known, but concentration of thand u-rich heavy minerals in a beach setting is thought to have been the dominant process. conclusions the present study has shown that member f-iii of the fjerritslev formation in the anholt borehole is of pliensbachian–toarcian age, member f-iv is of aalenian age and the haldager sand formation is of latest aalenian – bajocian/bathonian age. the sediments of the fjerritslev formation are mostly marine clays with varying proportions of storm sand laminae that generally increase up-section in abundance, thickness and grain-size. the haldager sand formation consists of an alteration of non-marine sand, clay and a few lignite layers. the sand layers are dominated by quartz and muscovite and, in the haldager sand formation, authigenic kaolinite and lignite clasts. the clays are dominated by kaolinite and illite with subordinate smectite, increasing in proportion down-section, and muscovite. in the lowermost haldager sand formation, the presence of heavy mineral-enriched fine-grained sand layers is inferred from geochemical anomalies. siderite and pyrite are commonly present as concretions, whereas calcite is only found in one concretion. the sand layers are unconsolidated and the clays still plastic. the diagenetic maturation of the organic material indicates that the sediments have been buried to a depth of about 1100–1300 m, and later uplifted to their present position, probably partly during late cretaceous – paleocene inversion and partly during the post-eocene regional uplift of parts of scandinavia. the detrital kaolinite exhibits a greater thermal stability (increasing down-section) than that expected in kaolinite at depths between 100 and 300 m. it is believed that the increased thermal stability was reached during progressive diagenesis, resulting from burial to a depth of 1200–1400 m. the progressive diagenesis did not reach the stage of complete illitisation of the original detrital smectite minerals. the lowermost part of the section, member f-iii of the fjerritslev formation, was deposited in a storminfluenced, marine, well-oxygenated environment. the influence of storms decreased somewhat in the latest pliensbachian and earliest toarcian, and the bottom conditions became more oxygen-deficient. throughout the remaining part of the toarcian, anoxic conditions prevailed and the environment shallowed progressively, probably due to localised, marginal high sedimentation rates despite the regional relative sea-level rise. deltaic environments dominated in latest toarcian – aalenian times, represented by the uppermost part of member f-iii and member f-iv of the fjerritslev formation. the bajocian/bathonian haldager sand formation is interpreted as a delta-plain deposit. acknowledgements we thank henrik friis for critically reading the manuscript, jette gissel nielsen who processed the foraminiferal samples and drafted some of the illustrations, bitten larsen and yvonne desezar for processing the palynological samples, and ulla bjerring, jette grejs pedersen, bente winsløv jensen and terese k. rasmussen for performing the sediment-analysis. lis olesen is thanked for drafting assistance. this work was supported by the danish natural science research council (geokat and magkat projects and an ole rømer stipendium to m.-s.s.) and the 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(eds): paleoecology, biostratigraphy, paleoceanography and taxonomy of agglutinated foraminifera, 621–657. dordrecht: kluwer academic publishers. nagy, j., gradstein, f.m., gibling, m.r. & thomas, f.c. 1995: foraminiferal stratigraphy and paleoenvironments of late jurassic to early cretaceous deposits in thakkhola, nepal. micropaleontology 41, 143–170. nielsen, b.l. & friis, h. 1985: diagenesis of middle jurassic haldager formation sandstone in the danish subbasin, north jutland. bulletin of the geological society of denmark 33, 273–285. nielsen, o.b. 1979: undersøgelse af jurassiske sedimenter i centralgraben, nordsøen, boring u-1, 16 pp. unpublished report, geologisk institut, aarhus universitet, danmark. nielsen, o.b. 1992: lithologi, lithostratigrafi og aflejringsmiljø i anholtboringen. dansk geologisk forening årsskrift for 1990–91, 67–72. nielsen, o.b., cremer, m., stein, r., thiébault, f. & zimmerman, 608 h. 1989: analysis of sedimentary facies, clay mineralogy, and geochemistry of the paleogene sediments of site 647, labrador sea. in: srivastava, s.p. et al. (eds): proceedings of the ocean drilling program, scientific results 105, 101–110. nielsen, s. 1985: aflejringsmiljø og lerdiagenese i det danske subbassin (øvre trias – nedre kridt) – belyst ved undersøgelse af cuttings fra dansk nordsø f-1 og boring 11/10-1, 154 pp. unpublished cand. scient. thesis, aarhus universitet, danmark. norling, e. 1972: jurassic stratigraphy and foraminifera of western scania, southern sweden. sveriges geologiska undersökning serie ca 47, 120 pp. norton, f.h. 1939: identification of clay minerals by d.t.a. journal of the american ceramic society 22, 54–63. nørvang, a. 1957: the foraminifera of the lias series in jutland, denmark. meddelelser fra dansk geologisk forening 13, 275–413. ogg, j.g., steiner, m.b., oloriz, f. & tavera, j.m. 1984: jurassic magnetostratigraphy, 1: kimmeridgian–tithonian of sierra gorda and carcabuey, southern spain. earth and planetary science letters 71, 147–162. pedersen, g.k. 1986: changes in the bivalve assemblage of an early jurassic mudstone sequence (the fjerritslev formation in the gassum 1 well, denmark). palaeogeography, palaeoclimatology, palaeoecology 53, 139–168. petersen, h.i., nielsen, l.h., bidstrup, t. & thomsen, e. 2003: burial depth and post-early cretaceous uplift of lower–middle jurassic strata in the fennoscandian border zone based on organic maturity. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 611–630 (this volume). pevear, d.r., williams, v.e. & mustoe, g. 1980: kaolinite, smectite and k-rectorite in bentonites: relation to coal rank at tulameen, british columbia. clays and clay minerals 28, 241–254. poulsen, n.e., gudmundson, l., hansen, j.m. & husfelt, y. 1990: palynological preparation techniques, a new macerationtankmethod and other modifications. danmarks geologiske undersøgelse serie c 10, 22 pp. press, f. & siever, r. 1982: earth, 3rd edition, 613 pp. san francisco: w.h. freeman & co. riding, j.b. & thomas, j.e. 1992: dinoflagellate cysts of the jurassic system. in: powell, a.j. (ed.): a stratigraphic index of dinoflagellate cysts, 7–97. british micropalaeontological society publication series. london: chapman & hall. schmidt, b.j. 1985a: clay mineral investigations of the rhaetic – jurassic – lower cretaceous sediments of the børglum-1 and uglev-1 wells, denmark. bulletin of the geological society of denmark 34, 97–110. schmidt, b.j. 1985b: a coal petrographic source rock evaluation of the rhaetic – jurassic – lower cretaceous sediments of the børglum-1 and uglev-1 wells, denmark. bulletin of the geological society of denmark 33, 239–252. schmidt, b.j. 1989: maturity and source rock evaluation of the mesozoic sequence in some danish off-shore wells outside the central trough, 429 pp. unpublished ph.d. thesis, aarhus university, denmark. schmidt, b.j. 1990: use of vitrinite reflectance in basin studies. in: balling, n. et al. (eds): proceedings of basin workshop, aarhus 1989. geoskrifter 35, 117–119. schulz, e. 1967: sporenpaläontologische untersuchungen zur rhät–lias–grenze in thüringen und der altmark. geologie 11(3), 308–320. seidenkrantz, m.-s., koppelhus, e.b. & ravn-sørensen, h. 1993: biostratigraphy and palaeoenvironmental analysis of a lower to middle jurassic succession on anholt, denmark. journal of micropalaeontology 12, 201–218. sorgenfrei, t. & buch, a. 1964. deep tests in denmark, 1935–1959. danmarks geologiske undersøgelse iii. række 36, 146 pp. środoń, j. & eberl, d.d. 1984: illite. in: bailey, s.w. (ed.): micas. reviews in mineralogy 13, 495–544. washington dc: mineralogical society of america. steiner, m.b., ogg, j.g., melendez, g. & sequeiros, l. 1985: jurassic magnetostratigraphy, 2: middle–late oxfordian of aguilon, iberian cordillera, northern spain. earth and planetary science letters 76, 151–166. steiner, m.b., ogg, j.g. & sandoval, j. 1987: jurassic magnetostratigraphy, 3: bathonian–bajocian of carcabuey, sierra harana and campillo de arenas (subbetic cordillera, southern spain). earth and planetary science letters 82, 357–372. strømnes, e. 1991: kvantifisering av tertiær oppløft og erosion i en travers over egersundbassenget, 71 pp. unpublished thesis, norges tekniske høgskole, trondheim, norge. teichmüller, m. & teichmüller, r. 1979: diagenesis of coal (coalification). in: larsen, g. & chilingar, r.v. (eds): diagenesis of sediments and sedimentary rocks, 207–246. amsterdam: elsevier. ter-borch, n. 1987: structural map of the top chalk group, 1:500 000. hørsholm, denmark: skovog naturstyrelsen, havbundsundersøgelsen og dansk olie& gasproduktion a/s. (also distributed as ter-borch, n. 1991: geologisk kort over danmark, 1:500 000. kalkoverfladens struktur. danmarks geologiske undersøgelse kortserie 7, 4 pp, 1 map.) thomsen, e., damtoft, k. & andersen, c. 1987: hydrocarbon plays in denmark outside the central trough. in: brooks, j. & glennie, k.w. (eds): petroleum geology of north west europe, 375–388. london: graham & trotman. tissot, b.p., pelet, r. & ungerer, p. 1987: thermal history of sedimentary basins, maturation indices and kinetics of oil and gas generation. american association of petroleum geologists bulletin 71, 1445–1466. waples, d.w. 1985: geochemistry in petroleum exploration, 232 pp. boston: international human resources development corporation. woollam, r. & riding, j.b. 1983: dinoflagellate cyst zonation of the english jurassic. institute of geological sciences report 83/2, 42 pp. london: her majesty’s stationery office. 609 manuscript received 2 september 1994; revision accepted 13 may 1997. geological survey of denmark and greenland bulletin 41, 2018, 25-28 25 glaciotectonic deformations often result in a high degree of variability, including glaciotectonic and sedimentary variability. redeposition of sediments during deformation increases the variability. ground-penetrating radar (gpr) has proven to be a good method to determine sedimentary structures in glaciofluvial deposits (olsen & andreasen 1994; van overmeeren 1998) as well as glaciotectonic structures (busby & merrit 1999; overgaard & jakobsen 2001). reflection facies analysis (radar facies) is a useful tool in the characterisation and interpretation of deformed sediments (van overmeeren 1998; jakobsen & overgaard 2002; lerche et al. 2014). a gpr survey was carried out at jyderup skov in odsherred in north-west sjælland (fig. 1). the presence of parallel ridges in the area indicates glaciotectonic deformation. the aim of the gpr study was to map the interior of the ridge complex and to interpret the genesis of the ridges. geological setting the morphology of odsherred in the north-western part of sjælland is dominated by three large arc-shaped ice-push ridges (fig. 1a). the ice-push ridges (arcs) were formed in the late weichselian during the bælthav readvance and glaciofluvial deposits related to the ridge formation have been dated to be c. 17 000 years (houmark-nielsen 2008). the arcs cut each other and they were formed by three ice readvances of ice-lobes situated east of them. each arc is a polymorphological landscape built-up of several landscape types. data from boreholes show that the arcs contain dislocated paleocene clay and weichselian marine deposits indicating that the interior of the arcs is affected by glaciotectonic deformation. the investigated area is located in the western part of the vig arc (fig. 1b), which is situated in the centre of three arcs in odsherred with the reclaimed sidinge fjord to the east forming a central depression. on the stoss side towards sidinge fjord is a terrain of smooth ground moraine. the upper part of the vig arc is characterised by hummocky moraine and the western part consists of elongated ice-marginal moraines. larger flat-topped kames occur within the hummocky moraine landscape and west of the ice-marginal moraine glaciofluvial deposits form an outwash plain. a number of elongated parallel ridges occur in the eastern part of the outwash plain (fig. 1c). a raw-material investigation shows that at least the upper 10 m consist of sand and gravel (region sjælland 2011), and borehole information shows the presence of sandy till. sedimentological and glaciotectonic interpretation of georadar data from the margin of the vig ice-push ridge, nw sjælland, denmark cecilie skovsø andersen and peter roll jakobsen 50 km sjælland odsherred jylland sweden germany c line05 1 km 43 m 1 m10 km c vig arc sidinge fjord 121 m –8 m ba b sejerø bugt fig. 1. a: map of denmark showing the location of odsherred. b: terrain model of odsherred, nw sjælland, with index map. c: terrain model of the investigated area showing the location of the georadar line line05 over the ice-margin ridges. © 2018 geus. geological survey of denmark and greenland bulletin 41, 25–28. open access: www.geus.dk/bulletin http://www.geus.dk/bulletin 2626 georadar survey and processing parameters five gpr lines were made perpendicular to the ridges and one was recorded parallel to the overall strike of the ridges. the occurrence of forest roads and paths in jyderup skov determined the location of the lines. the data were acquired in october 2015 using a 250 mhz sensors & software inc. georadar. due to the forest setting shielded antennae on a skid plate were used with a c. 40 cm offset. the traces were sampled with a step size of c. 5 cm and then stacked by a factor of 8. the line line05 provides the best image of the internal structure of the ridges (figs 1b, 2). processing of the raw line05 was carried out using pulseekko software by sensors & software inc. in the following steps: first, a dewow filter with an operational length of 1 pulse width was applied, followed by a stolt migration with an estimated velocity of 0.133 mns–1, which is typical value for dry sand. an automatic gain control with a pulse width of 1 and a maximum scaling factor of 700 was then applied, before the final topographic corrections were added. the vertical resolution is c. 20 cm. georadar facies description and interpretation the line05 profile is 850 m long and cross-cuts 10 ridges excluding the moraine and the two ridges behind the icemarginal moraine of the vig arc (fig. 2). four radar facies are distinguished in the profile. radar facies 1 (fig. 2) is interpreted as glaciofluvial sand and gravel based on the relatively strong parallel to subparallel reflections with high contrast, which are moderately continuous to discontinuous. borehole data indicate that glaciofluvial sand and gravel are found to a depth of c. 20 m below the terrain surface. the reflections are primarily planar to wavy shaped, but antiforms and synforms are also found (8.5 m wide on average; figs 2, 3a). the antiforms have a slightly asymmetric appearance with the western flank being steeper 800 m 700 m 600 m 10 m 14 m 0 ns 50 ns ew radar facies 1 radar facies 2 radar facies 4radar facies 3 thrust fault 400 m 300 m 16 m 20 m 0 ns 50 ns fig. 3a fig. 3b 200 m 100 m 0 ns 50 ns 0 m 22 m 26 m inferred marker horizon 500 m fig. 2. interpretation of the ground-penetrating radar profile line05. the scale to the left shows metres above sealevel. the scale to the right shows twoway travel time in nanoseconds. the black line is an inferred marker horizon used for structural analysis. the vertical exaggeration is 1:4. the black boxes indicate details shown in fig. 3. 27 (typically c. 12°) than the eastern flank (typically c. 6°). the antiform has an interlimb angle of typically c. 160°. the reflectors show frequent displacements, which we interpret as thrust faults (figs 2, 3a). the offset is between 0.20 m and 4.80 m with an average of 1.36 m. the reflections dip 4.0° to 13.6° dominantly to the east and are generally steepening upwards in a gentle concave to steep convex dipping pattern. the facies is truncated by facies 2, 3 and 4 in a concordant and erosional way (fig. 2). we describe the detailed characteristics of the thrusts in the next section. in radar facies 2 (figs 2, 3b) the reflections are discontinuous and show a relatively low amplitude contrast. the reflection pattern is chaotic at c. 280 m to subparallel at c. 325 m (fig. 3b) and c. 75 m, where the reflections dip downwards at the ends of the facies. based on the facies characteristics, this facies is interpreted as a gravelly ablation till. facies 2 is found in patches and it drapes facies 1. radar facies 3 (figs 2, 3a) consists of steepening upward (on average 10°) moderately continuous to discontinuous and subparallel reflections, this facies is mainly found on the western slope of the ridges. the reflection pattern is similar to facies 1. based on the position of the facies and the reflection pattern, facies 1 is likely to be the source of the sediments of facies 3 that were redeposited as solifluction sediments. this interpretation is in good agreement with the primary location of the facies on the western flank of the ridges, where it is more inclined to slope failure. in radar facies 4 (fig. 2) the reflection pattern can be described as continuous, parallel and diverging with a high density of downlaps and onlaps in the crests between ridges. the facies drapes the other three facies, and in some areas, it truncates facies 1 in an erosional way. facies 4 has a strong relative amplitude contrast. based on the reflection pattern, facies 4 is interpreted as sand and gravel. the facies is typically thin and lacking on top of some of the ridges (figs 2, 3a). however, it is locally thick on the crests. the facies is interpreted as post-deformation deposits. structural geology the thrust-fault planes are commonly concave and convex rotating listric faults steepening upwards following and cross-cutting the bedding, with a dip of 2.5° to 21.9° (13.4° on average) mainly to the east. the westward dipping thrust faults are usually small and often back-thrusts from the more extensive eastward dipping thrusts. some of the thrusts are clustered in a complex splay pattern, where the main thrust jacks up with each splaying thrust. these thrust faults are interpreted as linked contractional, leading imbricate stacks, originating from the same root at a detachment. the displacement is greatest at the frontal thrust sheet. the leading imbricate stacks are primarily located beneath the ridges on the eastern flank, but they also occur on the western flanks in association with antiforms and synforms. fault-bend-folding style of thrusting with hanging-wall ramp-flat-ramps and footwall flat-ramp-flats is evident from the concave–convex shape of the bedding along the fault plane. fault propagation folding is seen in the imbricate stacks where the tip line terminates in the bedding. by drawing an inferred marker horizon (fig. 2) on line05, the shortening of the line, caused by thrusting and folding, is calculated to be 12% of its initial length before deformation. formation of the ridges the formation of the ridges can be described tectono-stratigraphically in three steps: pre-tectonic deposits: radar facies 1 was deposited as glaciofluvial sand and gravel on a pro-glacial outwash plain. during a synand post-depositional event, radar facies 1 was folded and thrust into gentle antiand synforms. the thrusts are dense on the eastern flank of the antiforms, and occur as thick, wide, complex imbricate fans. the slight asymmetry of the antiand synforms and the dip direction of the thrusts 20 m 26 m 20 m 26 m 0 ns 50 ns a 0 ns 50 ns b w e fig. 3. details from fig. 2 (same legend as fig. 2). a: ridge formed by thrusting, and subsequently smoothened by radar facies 3 and 4. b: relationship between radar facies 1, 2 and 4. 2828 and bedding indicate a direction of primary deformational stress from the east. tectonically penecontemporaneous deposits: radar facies 2 was deposited penecontemporanously as an ablation till. it occurs in patches and is only seen in one borehole (dgu no 190.196) c. 250 m south-west of line05, where the upper 2 m consist of sandy and gravelly till. radar facies 2 is recognised in three places along line05, at c. 75 m, c. 280 m and c. 325 m. at c. 325 m, the radar section cuts an elevated area, which does not have the distinct ridge morphology seen elsewhere. the flattened morphology is in good agreement with deposition of ablation till, in contrast to the deformed elongated ridges elsewhere in the area. at c. 75 m the radar section is within an area of former stagnant ice, where ablation till should be expected. post-tectonic deposition: after the deformation came to an end, erosion of radar facies 1 occurred, and some thrust sheets are clearly truncated by solifluction sediments (radar facies 3) and post-deformation deposits (radar facies 4). the small parallel ridges west of the large vig arc may have been formed in a single deformational event creating a thinskinned thrust-fault complex in a gravity-spreading environment (jakobsen & overgaard 2002; pedersen 2005) or as annual moraines formed at an oscillating ice-margin (krüger 1995; benediktsson et al. 2009). the presence of till deposits within the area with the small ridges indicates that the ice margin was situated in the area. the topographic relief of the ridges is pronounced and there is no indication of erosion of the top of the ridges. thus, the glacier has not moved past the individual ridges. we therefore suggest that the ridges were formed as ice-marginal push moraines created by seasonal advances and retreats of the ice margin. conclusions ground-penetrating radar was used to map the sedimentological and glaciotectonic structures of a series of parallel morine ridges of an area west of the vig arc in odsherred. the gpr profiles have a high resolution that allows detailed sedimentological and structural analyses. four radar facies are distinguished which represent different sedimentary environments and degrees of deformation and three tectonostratigraphic sequences are recognised: pre-tectonic, tectonically penecontemporaneous and post-tectonic deposits. the interior of the ridges is characterised by thrust faults and folds created by deformation from the east and the ridge morphology is clearly associated with the deformation structures. subsequently the ridge morphology has been smoothened by post-tectonic sedimentation. we suggest that ridges seen on the proximal part of the outwash plain, west of the vig arc, were formed as ice-push moraines by seasonal advances and retreats of the ice margin. references benediktsson, í.ö., ingólfsson, ó., schomacker, a. & kjær, k.h. 2009: formation of submarginal and proglacial end moraines: implications of ice-flow mechanism during the 1963-64 surge of brúarjökull, iceland. boreas 38, 440–457. busby, j.p. & merrit, j.w. 1999: quaternary deformation mapping with ground penetrating radar. journal of applied geophysiscs 41, 75–91. houmark-nielsen, m. 2008: testing osl failures against a regional weichselian glaciation chronology from southern scandinavia. boreas 37, 660–677. jakobsen, p.r. & overgaard, t. 2002: georadar facies and glaciotectonic structures in ice marginal deposits northwest zealand, denmark. quarternary science reviews 21, 917–927. krüger, j. 1995: origin, chronology and climatological significance of annual-moraine ridges at myrdaljökull, iceland. the holocene 5, 420–427. lerche, h., jakobsen, p.r. & pedersen, s.a.s. 2014: ribbed moraines formed during the retreat of the scandinavian ice sheet from eastern himmerland, ne jylland, denmark. geological survey of denmark and greenland bulletin 31, 39–42. pedersen, s.a.s. 2005: structural analysis of the rubjerg knude glaciotectonic complex, vendsyssel, northern denmark. geological survey of denmark and greenland bulletin 8, 192 pp. olsen, h. & andreasen, f. 1994: sedimentology and ground-penetrating radar characteristics of a pleistocene sandur deposit. sedimentary geology 99, 1–15. overgaard, t. & jakobsen, p.r. 2001: mapping of glaciotectonic deformation in an ice marginal environment with ground penetrating radar. journal of applied geophysics 47, 191–197. region sjælland 2011: råstofkortlægning. rapport 8, sand, grus, sten, jyderup, odsherred kommune, 38 pp. http://jupiter.geus.dk/rapportdb/ grundvandsrapport.seam?grundvandsrapportrapportid=90959 van overmeeren, r.a. 1998: radar facies of unconsolidated sediments in the netherlands: a radar stratigraphic interpretation method for hydrogeology. journal of applied geophysics 40, 1–18. authors’ addresses c.s.a., orbicon, linnés alle 2, 2630 taastrup, e-mail: cesa@orbicon.dk. p.r.j., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. http://jupiter.geus.dk/rapportdb/grundvandsrapport.seam?grundvandsrapportrapportid=90959 http://jupiter.geus.dk/rapportdb/grundvandsrapport.seam?grundvandsrapportrapportid=90959 mailto:cesa@orbicon.dk geological survey of denmark and greenland bulletin 31, 2014, 71-74 71© 2014 geus. geological survey of denmark and greenland bulletin 31, 71–74. open access: www.geus.dk/publications/bull integrating 3d photogeology with aeromagnetic data as a tool for base-metal exploration in east greenland anaïs brethes, pierpaolo guarnieri and thorkild m. rasmussen an 800 km long basin system developed along the east greenland margin since the late palaeozoic in which the jameson land basin forms the southern part of the system. along the margins of the jameson land basin there are occurrences of barite, copper, lead, zinc and silver, which are particularly abundant in the northern part of the basin’s eastern margin in the wegener halvø area (fig. 1). structures and stratigraphic architecture play important roles in the mineralisation distribution, so detailed mapping is essential. we used 3d photogeology combined with geophysical data to map the different stratigraphic units, faults and dykes in three dimensions. geological setting of jameson land basin the east greenland basin began to develop during the devonian following the caledonian orogeny. in the late carboniferous and early permian rotational block faulting took place (surlyk 1990). the faults were reactivated during the late permian and early triassic and synsedimentary faulting occurred at the same time (seidler 2000). the resulting fault blocks were peneplained before transgression and onset of sedimentation in the late permian. on wegener halvø the up to 300 m thick upper permian sequence unconformably overlies deformed and peneplained devonian sediments (fig. 1). the sequence is dominated by karstified dolomitic limestone of the karstryggen formation, by carbonate buildups which developed on palaeotopographic highs (the wegener halvø formation) and by bituminous dark shale of the upper permian ravnefjeld formation. these filled up the deepest and karstified parts of the previous formations and locally contain calcareous concretions and fossils derived from the carbonate buildups. the unconformable to conformable permian–triassic boundary is locally incised by submarine canyons (seidler 2000). laterally variable, the triassic stratigraphic architecture was controlled by tectonics and rapid asymmetric subsidence. this entire sequence is mainly composed of continental deposits with some shallow marine incursions. in the tertiary the jameson land basin was covered by flood basalt and intruded by a complex of sills and e–w to nw–se-trending dykes. the intrusions that are dated to the latest paleocene to earliest eocene (soper et al. 1976) affect the entire sequence of basin fill. in miocene time the northern part of the basin was uplifted more than 1 km (mathiesen et al. 2000; hansen et al. 2001). base-metal mineralisation on wegener halvø the first rock samples were collected in east greenland in 1822 and during the three-year expedition to east greenland from 1931 to 1934 copper-, leadand silver-bearing minerals were discovered on wegener halvø (fig. 1), in the upper permian carbonates of the wegener halvø formation. subsequent investigations by nordisk mineselskab a/s (nordmine) led to the discovery in 1968 of base-metal occurrences in the black shale of the ravnefjeld formation. min5 km greenland 22°45´2.5 km precambrian to lower permian upper permian triassic jurassic quaternary fault 22°30´ 71°35´ 71°40´ nordenskiöld bjerg buch bjerg quensel bjerg tait bjerg fle ming fjor d nath ors t fj ord per mdal devo ndal weg en er halvø ss jl ll fig. 1. simplified geological map of the wegener halvø area (modified from perch-nielsen et al. 1983). ss: scoresby sund. jl: jameson land. ll: liverpool land. contour interval 200 m. 7272 eral exploration ceased because the ore was too low grade, but the oil industry became interested in the hydrocarbon source-rock potential of the upper permian shale. harpøth et al. (1986) reviewed more than a century of these sporadic field campaigns and since 2011 avannaa resources ltd has explored the area for base metals. mineral occurrences are particularly abundant along the eastern margin of the jameson land basin, especially in the wegener halvø area where they occur in three of the upper permian formations and three of the triassic formations. one of the most interesting ore-bearing formations is the limestone buildups of the wegener halvø formation. as described by harpøth et al. (1986), cu, pb and zn sulphides are scattered throughout the whole peninsula. they are also found as fine-grained material in inter-reef black-shale occurrences in the ravnefjeld formation. base-metal minerals are mainly found in the upper part of the carbonates and in the lower part of the shales (upper permian formations), but stratiform occurrences occur in triassic sediments. early triassic alluvial conglomerate and arkosic sandstone are mineralised with argentiferous chalcocite-covellite, and galena occurs as cement. according to soil sample geochemical analyses, the highest grades appear to be concentrated at the contact between two facies, in a trend perpendicular to the palaeocurrents (harpøth et al. 1986). in the late triassic playa-lake mudstone and sandstone, fine-grained cu sulphides are found over an area of 1000 km2. 3d photogeology the geological survey of denmark and greenland is equipped with a 3d stereoplotter to carry out multi-model photogrammetry that can be used to accurately map geological features from vertical and oblique aerial photographs. the instrument is equipped with two polarised monitors set one above the other with an angle of 110° separated by a halfmirror (vosgerau et al. 2010). each polarised screen displays one photograph of an overlapping stereopair. the half-mirror allows the viewer to see two photographs as one image with 3d polarised glasses getting a depth perception of the geological features that, in turn, can be digitised as 3d polylines. the geologic features and the database are automatically imported into a geographic information system where they can be edited. the photographs used for the present work were taken during a combined hyperspectral and lidar survey in 2012 (thorning et al. in press). the spacing between the flight lines was c. 500 m and the pixel size in the centre of each image is 27 × 27 cm. on wegener halvø, an area of 100 km2 comprising devondal and nordenskiöld bjerg was mapped on a scale of 1:1000 (fig. 1). more than 5000 features were digitised, including bedding, stratigraphic boundaries as well as fractures and dykes (fig. 2). together with the digitised 3d polylines it is possible to calculate strata thicknesses and measure strike and dip of bedding of faults and dykes. the overall fig. 2. 3d view from the south towards nordenskiöld bjerg created with orthophotographs draped on a digital elevation model developed from lidar data. the geological features collected from aerial photographs are: faults and fractures (red lines) and triassic bedding (black lines). 400 m 73 stratigraphic succession is based on the published geological map (perch-nielsen et al. 1983). more than 900 faults, fractures and dykes were digitised in the wegener halvø area and the majority shows a n–s and nnw–sse-orientation. the structures are steep and occur as flower structures affecting the entire succession. the vertical offset is small and decreases from south to north along the strike of the faults that in some cases have a strike-slip component. integrated 3d photogeology and aeromagnetic data the dataset from the interpretation of the photographs has been combined with a new structural interpretation of airborne magnetic data from the aem greenland 1997 survey (rasmussen et al. 2001). different maps were obtained from the calculation of the magnetic gradient tensor elements to highlight structures in different directions (pedersen & rasmussen 1990). the total magnetic intensity anomalies of the structures of interest have small amplitudes (around 3 nt) that are close to the detection limit of c. 0.5 nt of an airborne survey. as differentiation works as a high-pass filter, the second vertical derivative of the total magnetic intensity was also calculated to emphasise the linear features. this processing provides a better dynamic range for visualising the structures than the measured field. to attenuate the short wavelength noise features generated by this differentiation, an upward continuation of the magnetic field to 100 m was applied. lineaments defined by negative magnetic anomalies and interpreted as major structures were digitised on the resulting map (fig. 3a). these features are n–s to nnw–sse-trending and several kilometres long and they fit with the faults in the geological map (fig. 1) and with the fault segments mapped with the 3d stereo-plotter (fig. 3b). the accurate fault segments mapped by photogeology were integrated with the magnetic lineaments highlighted at a larger scale. thus, an identified structure can be prolongated on the map using geophysical data and constrained in its precise location and dip using 3d photogeology. one of the n–s-trending faults mapped by 3d photogeology and visible on the magnetic data coincides with the so-called vimmelskaft lineament defined by pedersen (1997). this lineament corresponds to a zone where a n–s-trending fault and a n–s-running dyke cut through upper permian shales and karstified carbonates. important mineral occurrences with high base-metal content are found along this lineament and the concentration decreases away from the lineament. fig. 3. a: second vertical derivative of the total magnetic intensity (tmi) from the aem greenland 1997 survey (pixel size is 50 × 50 m) continued upward to 100 m and draped on the shaded elevation model. the negative magnetic lineaments are digitised and shown as black lines. b: correlation between magnetic data and the faults, fractures and bedding collected from aerial photographs. note the presence of eroded strata parallel to positive magnetic anomalies and faults and fractures related to negative anomalies. lineament 2nd vertical derivative tmi (nt/m2) –1.2 × 10–5 1.5 × 10–5 b a 71°28´ 22°48´ 4 km -0.3.10 -5 [nt.m -2] 0.5.10 -5 fleming fjord fm. gipsdalen fm. pingodal fm. faults and fractures beddings: fleming fjord fm gipsdalen fm pingo dal fm faults and fractures bedding: 0.5 × 10–5 –0.3 × 10–5 500 m 71°34´ 22°30´b 2nd vertical derivative tmi (nt/m2) 7474 the positive magnetic anomalies are not linear and appear to be associated with magnetic domains in the triassic formations (fig. 3b). the contours of these domains are parallel to the stratification collected from 3d photogeology (edges of eroded beds) and are mostly located on crests. concluding remarks using 3d photogeology we mapped a 100 km2 area with good exposures on a scale of 1:1000. formation boundaries and dip of strata are well constrained. new structures were mapped and the resulting dataset forms a good starting point for further studies of the stratigraphy or tectonic evolution using 3d modelling. the combination of 3d photogeology with new processing of the magnetic data has shown a good match between n–s-trending magnetic lineaments and n–s-trending normal steep faults. structures of this type were highlighted by pedersen (1997) for their possible role as pathways for mineralising fluids. the structures mapped during this study that affect the upper permian carbonates and shales, should be checked in the field. furthermore, although mineral occurrences found in the triassic formations are stratiform, remobilised sulphides are concentrated in cross-cutting fractures. the mapped faults and fractures that cut through triassic formations could also be interesting targets for exploration. detailed studies of structures and stratigraphic architecture are key points for mineral exploration. accurate mapping using 3d photogeology combined with geophysical data can be efficient tools for this, especially in areas with difficult access and excellent outcrops. references hansen, k., bergman, s.c. & henk, b. 2001: the jameson land basin (east greenland): a fission track study of the tectonic and thermal evolution in the cenozoic north atlantic spreading regime. tectonophysics 331, 307–339. harpøth, o., pedersen, j.l., schønwandt, h.k. & thomassen, b. 1986: the mineral occurrences of central east greenland. meddelelser om grønland, geoscience 17, 139 pp. mathiesen, a., bidstrup, t. & christiansen, f.g. 2000: denudation and uplift history of the jameson land basin, east greenland – constrained from maturity and apatite fission track data. global and planetary change 24, 275–301. pedersen, m. 1997: investigation of ore potential in black shales of the upper permian ravnefjeld formation in scoresby land and on traill ø, central east greenland. danmarks og grønlands geologiske undersøgelse rapport 1997/124, 17 pp. pedersen, l.b. & rasmussen, t.m. 1990: the gradient tensor of potential field anomalies. some implications on data collection and data processing of maps. geophysics 55, 1558–1566. perch-nielsen, k., henriksen n. & stemmerik, l. 1983: geological map of greenland, 1:100 000, fleming fjord, 71 ø.1 nord. copenhagen: geological survey of greenland. rasmussen, t.m., thorning, l., stemp, r.w., jørgensen, m.s. & schjøth, f. 2001: aem greenland 1994–1998 – summary report. danmarks og grønlands geologiske undersøgelse rapport 2001/58, 46 pp. seidler, l. 2000: incised submarine canyons governing new evidence of early triassic rifting in east greenland. palaeogeography, palaeoclimatology, palaeoecology 161, 267–293. soper, n.j., downie, c., higgins, a.c. & costa, l.i. 1976: biostratigraphic ages of tertiary basalts on the east greenland continental margin and their relationship to plate separation in the northeast atlantic. earth and planetary science letters 32, 149–157. surlyk, f. 1990: timing, style and sedimentary evolution of late palaeozoic – mesozoic extensional basins of east greenland. in: hardman, r.f.p. & brooks, j. (eds): tectonic events responsible for britain’s oil and gas reserves. geological society, special publication (london) 55, 107–125. thorning, l., christensen, n.n., olsen, s., riisager, p., sørensen, l.l., sørensen, e.v. & tukiainen, t. in press: high resolution airborne hyperspectral imaging spectroscopy in central east greenland 2012 – data acquisition and pre-processing. danmarks og grønlands geologiske undersøgelse rapport. vosgerau h., guarnieri p., weibel r., larsen m., dennehy, c., sørensen, e.v. & knudsen, c. 2010: study of a palaeogene intrabasaltic sedimentary unit in southern east greenland: from 3-d photogeology to micropetrography. geological survey of denmark and greenland bulletin 20, 75–78. authors’ addresses a.b. & p.g., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: aib@geus.dk t.m.r., division of geosciences and environmental engineering, luleå university of technology, s-971 87 luleå, sweden. geological survey of denmark and greenland bulletin 4, 2003, pp 57-60 57 prior to the 1990s only few geological investigations of the seabed and the shallow geology around the faroe islands had been undertaken (waagstein & rasmussen 1975; nielsen et al. 1981). however, in the 1990s marine geological and in particular seismic investigations were markedly intensified. since 1993 several studies on the structure of the faroe islands margin and seafloor processes have been funded by the european union, namely the enam (european north atlantic margin) project i and ii (1993–1999) and the stratagem (stratigraphy of the glaciated european margin) project (2000–2003), and these have provided significant new information on the mechanisms shaping the faroe islands margin (e.g. boldreel et al. 1998; kuijpers et al. 1998a; nielsen & van weering 1998; van weering et al. 1998). due to the expertise and regional geological knowledge obtained during these projects the geological survey of denmark and greenland (geus) became involved in socalled ‘geohazard’ seabed studies of the faroe–shetland channel in 1997. these investigations were financed by the petroleum industry that had begun to show significant interest in exploration of the faroe–shetland channel area. the studies focused on possible natural risks that would affect geohazard studies offshore the faroe islands: slope instability, bottom currents and sub-seabed sediment mobilisation tove nielsen and antoon kuijpers fig. 1. overview of the faroe platform area with locations of major mass flow deposits, the pathway of high-energy norwegian sea overflow water (nsow), and location of the mud diapirs at the northern entrance of the faroe–shetland channel. fs, foib slide; gr, gem raft; sandf, sandoy fan; sands, sandoy slump; sudf, su›uroy fan; suds, su›uroy slump. geological survey of denmark and greenland bulletin 4, 57–60 (2004) © geus, 2004 submarine structures, such as slope instability and strong bottom currents, and included both shallow seismic data acquisition and sediment core analyses. most of the work at sea was undertaken with the russian research vessel prof. logachev, and carried out within the framework of the international, unesco-supported ‘training-through-research’ (ttr) programme co-ordinated by moscow state university, russia. since 1997, more than three million dkk have been granted for various projects and this work has been documented in 14 classified reports. this paper presents some of the main results from these ‘geohazard’ studies, in particular with respect to the sediment instability affecting the western flank of the faroe–shetland channel, the occurrence of very strong bottom currents in the channel, and the newly discovered mud diapirs at the northern entrance of the channel (fig. 1). material and methods seismic data acquisition was carried out with a 100 khz airgun and a 6-channel streamer to map subbottom structures. seismic profiling was carried out in combination with a 10 khz long-range (2 × 6 km) side-scan sonar for obtaining information on seabed surface sediment and topography. in selected areas a higher resolution of the seabed features was necessary, and a deep-towed side-scan sonar was deployed, operating at 30 or 100 khz with ranges of 2 × 1000 m and 2 × 350 m, respectively. the latter device was also equipped with a 5 khz subbottom profiler, whereas during all survey activities another, hull-mounted subbottom profiler was routinely operated. bottom samples were retrieved with a 6-m gravity corer, a box corer, and a large video-controlled grab. in addition, underwater video was deployed. after retrieval of the sediment cores, the cores were described and magnetic susceptibility measurements were carried out on board. selected samples were investigated using a microscope to determine mineral and microfossil content. after the cruise more extensive core studies were made, and sediments were dated using the ams 14c method. slope instability prior to of the ‘geohazard’ studies, mass flow deposits had not been reported from the western flank of the faroe–shetland channel. in contrast, a major slide complex was known to extend over most of the north-eastern faroe islands margin (fig. 1). seismic studies carried out in the latter area during the enam project demonstrated that large-scale slumping and sliding had affected the middle and lower slopes below 1500 m water depth since the miocene (nielsen & van weering 1998; van weering et al. 1998). high-resolution side-scan sonar surveying in the area downslope of the main, c. 300 m high headwall, where water depth is about 2300 m, demonstrated the presence of a large number of downslopetrending tracks on a low slope gradient, locally displaying cross patterns, and occasionally a markedly irregular pattern (fig. 2). at the termination of the tracks, outrunner blocks of sediment were observed, up to 18 m high, and with a maximum length of 70 m. some of the blocks were found at a distance of up to 25 km from the initial mass flow terminus at the main headwall. the sub-bottom profiles in the trail-mark area indicate that most of the tracks have been filled with transparent sediment acoustically comparable to the holocene hemipelagic surface unit, and thus may have an age older than holocene. at the start of the studies of the western flank of the faroe–shetland channel, it soon became evident that slope instability and associated mass flow had also occurred in this area. seismic evidence (fig. 3) clearly shows that these processes have taken place repeatedly since late pliocene time. within this context, it should be noted that no evidence has been found for any major mass-wasting activity having occurred subsequent to the early holocene sea level rise (kuijpers et al. 2001); i.e. during the past c. 7000 years the faroe islands margin appears to have been generally stable. high-energy bottom current environments export of deep waters formed in the north atlantic occurs via two major gateways: one between greenland and iceland and one between iceland and scotland (e.g. dickson et al. 1990). sediment core studies have demonstrated that these 58 fig. 2. deep-tow side-scan sonar record and sub-bottom profile from the trail mark area downslope of the main slump scar in the mass flow area north-east of the faroe islands (see fig. 1). the sonograph shows a large number of outrunner block tracks of varying width, with several blocks (circles) at the ends of their respective tracks. crossing slide paths are also observed. arrow indicates markedly irregular pattern of some tracks. from kuijpers et al. (2001). overflow currents were generally reduced or ceased during cold (stadial) climate periods, relative to interstadial and particularly interglacial conditions (e.g. kuijpers et al. 1998b). noteworthy in this context is the recent observation in the faroe–shetland channel of a decreasing overflow since 1950 (hansen et al. 2001). for the purpose of providing information on the high-energy overflow current environments along the faroe–shetland gateway, an inventory of current-induced bedforms detected by side-scan sonar was undertaken, which revealed the flow path where near-bottom current speed reaches around 1.0 m/s (kuijpers et al. 2002). for comparison, supplementary information from actual current meter measurements has been added in order to determine whether the bedforms recorded (fig. 4) could be relict features, or can be considered to be in equilibrium with the recent current regime. our knowledge of overflow processes, which were previously based only on information from current meter stations and ship-borne hydrographic sections, has thus been extended and a regional overview of the areas most intensively influenced by the overflow currents has been obtained. sub-seabed sediment mobilisation submarine mud volcanoes, or diapirs, can range in size between 0.5 and 800 m high. two main mechanisms are considered to lead to the formation of mud diapirism, i.e. high sedimentation rates and/or lateral tectonic compression. both mechanisms can result in over-pressure of a mobile sediment layer at sub-bottom depth. in the mid-1990s mound features were observed immediately east of the fugloy ridge (see fig. 1) by the british geological survey (bgs), and were reported as possible cold-water coral mounds. further high59 fig. 3. single-channel airgun profile from the western flank and basin of the faroe–shetland channel and a schematic interpretation (bottom) showing the presence of a large mass flow unit of presumably late pliocene – early pleistocene age. fig. 4. deep-tow side-scan sonar record of norwegian sea overflow water (nsow)-induced sandwaves at the southern end of the faroe–shetland channel. water depth is 1100–1200 m. from kuijpers et al. (2002). resolution seismic work by bgs and the royal netherlands institute for sea research (nioz) revealed, however, that the mounds were most likely mud-diapirs, an interpretation later supported by a tobi side-scan sonar survey carried out by the southampton oceanography centre (soc). during the 2002 ttr-cruise with r/v prof. logachev geus made a detailed study of the mounds at the northern entrance of the faroe–shetland channel (fig. 5). the results of this work confirm that the mound structures can be classified as mud diapirs originating from subsurface sediment mobilisation. this sediment mobilisation is probably due to the excessive load of dense, glacigenic sediments of the north sea fan deposited on top of low-density (miocene) diatomaceous ooze. several stages of maturity have been observed: (1) an initial stage where the diapirs do not pierce the seabed, (2) a young (up to 50 m high) stage displaying a marked relief, and (3) an up to 100 m high, mature stage where the diapirs have a smoother appearance. preliminary results from ams 14c dating of sediment cores collected from the diapirs suggest an episode of major activation of the diapirs around the time of the last glacial maximum (lgm). acknowledgements the studies were supported by the faroese offshore consortium gem (now foib) and the european enam-ii and stratagem projects. the contributions from colleagues at geus, nioz, bgs and the ttr-programme are gratefully acknowledged. references boldreel, l.o. & andersen, m.s. 1993: late paleocene to miocene compression in the faeroe–rockall area. in: parker, j.r. (ed.): petroleum geology of northwest europe: proceedings of the 4th conference, 1025–1034. london: the geological society. boldreel, l.o., andersen, m.s. & kuijpers, a. 1998: neogene seismic facies and deep-water gateways in the faeroe bank area, ne atlantic. marine geology 152, 129–140. dickson, r.r., gmitrowicz, e.m. & watson, a.j. 1990: deep water renewal in the northern north atlantic. nature 344, 848–850. hansen, b., turell, w.r. & østerhus, s. 2001: decreasing overflow from the nordic seas into the atlantic ocean through the faroe bank channel since 1950. nature 411, 927–930. kuijpers, a., andersen, m.s., kenyon, n.h., kunzendorf, h. & van weering, t.c.e. 1998a: quaternary sedimentation and norwegian sea overflow pathways around bill bailey bank, northeastern atlantic. marine geology 152, 101–127. kuijpers, a., troelstra, s.r., wisse, m., heier nielsen, s. & van weering, t.c.e. 1998b: norwegian sea overflow variability and ne atlantic surface hydrography during the past 150,000 years. marine geology 152, 75–99. kuijpers, a., nielsen, t., akhmetzhanov, a., de haas, h., kenyon, n.h. & van weering, t.c.e. 2001: late quaternary slope instability on the faeroe margin: mass flow features and timing of events. geo-marine letters 20, 149–159. kuijpers, a., hansen, b., hühnerbach, v., larsen, b., nielsen, t. & werner, f. 2002: norwegian sea overflow through the faroe–shetland gateway as documented by its bedforms. marine geology 188, 147–164. nielsen, t. & van weering, t.c.e. 1998: seismic stratigraphy and sedimentary processes at the norwegian sea margin northeast of the faeroe islands. marine geology 152, 141–157. nielsen, t.h., waagstein, r., rasmussen, j. & larsen, b. 1981: marine seismic investigation of the shelf around the faeroe islands. danmarks geologiske undersøgelse årbog 1981, 101–109. van weering, t.c.e., nielsen, t., kenyon, n.h., akentieva, k. & kuijpers, a. 1998: sediments and sedimentation at the ne faeroe continental margin; contourites and large-scale sliding. marine geology 152, 159–176. waagstein, r. & rasmussen, j. 1975: glacial erratics from the seafloor south-east of the faeroe islands and the limit of glaciation. annales societatis scientarum faeroensis 23, 101–119. 60 fig. 5. mosaic image of deep-tow side-scan sonar records of the mud diapir area at the northern entrance of the faroe–shetland channel. water depth is 1600–1700 m. from nielsen et al. (2002). authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: tni@geus.dk << /ascii85encodepages false /allowtransparency false /autopositionepsfiles true /autorotatepages /all /binding /left /calgrayprofile (dot gain 20%) /calrgbprofile (srgb iec61966-2.1) /calcmykprofile (u.s. web coated \050swop\051 v2) /srgbprofile (srgb iec61966-2.1) /cannotembedfontpolicy /warning /compatibilitylevel 1.4 /compressobjects /tags /compresspages true /convertimagestoindexed true /passthroughjpegimages true /createjdffile false /createjobticket false /defaultrenderingintent /default /detectblends true /colorconversionstrategy /leavecolorunchanged /dothumbnails false /embedallfonts true /embedjoboptions true /dscreportinglevel 0 /syntheticboldness 1.00 /emitdscwarnings false /endpage -1 /imagememory 1048576 /lockdistillerparams false /maxsubsetpct 100 /optimize true /opm 1 /parsedsccomments true /parsedsccommentsfordocinfo true 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<feff4f7f752890194e9b8a2d5b9a5efa7acb76840020005000440046002065874ef65305542b8f039ad876845f7150cf89e367905ea6ff0c4fbf65bc63d066075217537054c18cea3002005000440046002065874ef653ef4ee54f7f75280020004100630072006f0062006100740020548c002000520065006100640065007200200035002e0030002053ca66f465b07248672c4f86958b555f3002> >> >> setdistillerparams << /hwresolution [2400 2400] /pagesize [595.000 842.000] >> setpagedevice geological survey of denmark and greenland bulletin 28, 2013, 17-20 17 a baltic ice lake lowstand of latest allerød age in the arkona basin, southern baltic sea ole bennike and jørn bo jensen after the last deglaciation, the baltic sea underwent a complex salinity history and dynamic shore-level development with several lacustrine and marine stages: the baltic ice lake, the yoldia sea, the ancylus lake and the littorina sea (björck 1995). in connection with shallow seismic profiling in the south-western baltic sea, two marked and widespread erosional unconformities have been identified (jensen et al. 1997, 1999; lemke et al. 1998; larsen 2004). the older unconformity occurs within sediments deposited in the baltic ice lake, whereas the younger one separates baltic ice lake sediments from holocene lake and mire deposits. the latter unconformity is dated to the transition between the younger dryas and the holocene, corresponding to c. 11.7 cal. ka bp and formed due to a sudden drop in the level of the icedammed baltic ice lake of around 25 m, caused by ice recession from mt. billingen in south central sweden. the age of the older erosional unconformity is poorly constrained. however, it has been suggested that the level of the baltic ice lake also experienced a sudden drop during the allerød chronozone. the temperature increased during the allerød, and it is possible that the margin of the fennoscandian ice sheet also receded from the mt. billingen area at this time. if so, the older erosional unconformity may have formed in connection with an early drainage of the baltic ice lake. the question about such an early lake-level fall was discussed by björck (1995), who listed a number of arguments for and against it. one of the arguments for drainage is that evidence of a rapid lake-level fall is seen in both south-eastern sweden and in other baltic countries. a study in southwestern sweden indicates a high discharge of freshwater, and data from south central sweden indicate a significant glacial recession west of mt. billingen during the allerød, which could have led to drainage of the baltic ice lake. it has been suggested that drainage happened around 13 cal. ka ago (uścinowicz 2006; andrén et al. 2011). in connection with sediment coring in the arkona basin, firm evidence of an early lowstand was for the first time identified in a sediment core. an organic-rich sediment which can be referred to this lowstand was found, and material was submitted for 14c dating. the aim of this paper is to report on the age and its implications. the location of the coring site in the arkona basin is shown in fig. 1. the basin is up to 49 m deep and the surface of the till is found at a depth of 30 to 70 m (moros et al. 2002). material and methods sediment coring was carried out using a 10 cm diameter vibrocorer with a 6 m long steel tube. normally, coring positions are selected from high-resolution, shallow seismic profiles, but in the deeper parts of the arkona basin, gas-bearing, organic-rich, holocene marine sediments are widespread and hamper the use of seismic data. hence the core described here was collected at a site with poor seismic data. the core was sampled at a water depth of 39.5 m, at 54°45.005́ n, 13°45.876́ e. the core was collected c. 5 km ene of the northern end of a seismic profile published by lemke et al. (1998), where the unconformity is clearly seen. the lower unconformity is also seen on seismic data collected in the region during the baltic pipe survey (larsen 2004). we collected two cores at the site, designated 258000-1 and 258000-2. the first core was collected in plastic foil, split lengthwise in the ship laboratory and described and subsampled for palaeoecological analysis. the second core was collected in a pvc tube and cut into 1 m sections for storage. however, the second core penetrated somewhat deeper than the first core, and in the core catcher, clay was found with abundant plant remains underlain by clay-rich, mediumgrained sand. ten 1 kg samples from core 258000-1 and from © 2013 geus. geological survey of denmark and greenland bulletin 28, 17–20. open access: www.geus.dk/publications/bull 14°e 14°e 55°n 55° 54°30´n 13°e 13°e 55–60 50–55 45–50 40–45 35–40 75–80 70–75 65–70 60–65 30–35 25–30 20–25 15–20 10–15 5–10 depth (m) 20 km 0–5 54°30´n 258000 skåne bornholm møn rügen fig. 1. bathymetrical map of the arkona basin. the red dot shows the location of the studied sediment core. the core site is also shown on fig. 5. 1818 the core catcher of 258000-2 were brought to the geological survey of denmark and greenland, where the samples were wet sieved shortly after the cruise. fruits of cladium mariscus (a reed plant) were dried and shortly after submitted for accelerator mass spectrometry (ams) radiocarbon age determination at the leibniz laboratory for radiometric dating and isotope research in kiel, germany. results and discussion core 258000-1 consists of olive-grey mud from the core top down to a depth of 412 cm (fig. 2). some shells of macoma balthica and mytilus edulis were noted, which shows that the mud is marine and of holocene age. from 412 to 473 cm, clay without carbonate is found; this succession was divided into three units according to variations in colour and texture. shells and head shields of chydoridae (cladocerans, water fleas) are common, head capsules of larvae of chironomidae (non-biting midges) are present but rare, egg cocoons of the fish leach piscicola geometra were found in two samples, and statoblasts of the bryozoan cristatella mucedo in four samples (fig. 3; table 1). these invertebrate remains show that the clay was deposited in a lake. the presence of rare remains of betula sect. albae (tree birch) and pinus sylvetris (pine) indicates an early holocene age. the lower part was rich in radicells (tiny roots) of reed plants, which show that the shoreline was not far away from the coring site. the clay was probably deposited in the ancylus lake and perhaps in the yoldia sea. d ep th b el ow c or e to p (c m ) holocene marine mud indistinct layering holocene marine mud homogeneous fineto medium-grained sand late glacial clay late glacial clay with plant remains late glacial sand holocene lacustrine clay 0 100 200 300 400 500 fig. 2. composite lithological log of core 258000 from the arkona basin. the two lower units were only found in core 258000-2. the core numbers refer to the system used at the department of marine geology at the institute for baltic sea research in warnemünde. fig. 3. macrofossil diagram of the lower parts of cores 258000-1 and 258000-2 from the arkona basin. the hollow bars show remains not counted. 2 1 50 10 100 20 1 200 1 5 10 400 420 440 460 480 500 520 540 560 580 600 d ep th (c m ) be tu la se ct . a lba e sp . pin us sy lve str is ph ra gm ite s a us tra lis cl ad iu m m ar isc us ca re x s p. co m ar um p al us tre ci cu ta vi ro sa ra di ce lls c ha ra ce ae in de t. a m bl ys te gia ce ae in de t. m en ya nt he s t rif oli at a n ym ph ae a sp . pis cic ola ge om et ra c lad oc er a c hi ro no m id ae in de t. cr ist at ell a m uc ed o pi sc es in de t. sc irp us la cu str is rare common abundant terrestrial telmatic lacustrine li th ol og y 19 the clay is underlain by 2 cm of fineto medium-grained grey sand. this thin sand layer may have been deposited at or after the final drainage of the baltic ice lake. below the sand layer and down to the bottom of the core at 570 cm, grey clay is found. this unit is rich in carbonate and barren of fossils and it is interpreted as late-glacial clay deposited in the baltic ice lake during the younger dryas. from the deeper core 258000-2, a sample rich in plant remains was analysed for macrofossils. the plant remains were dominated by fruits and seeds of the telmatic plants menyanthes trifoliata and fruits of carex, mainly carex vesicaria, cladium mariscus, scirpus lacustris, comarum palustris and cicuta virosa (fig. 3). limnic plants and animals were represented by a seed of nymphaea sp., common stems of amblystegiaceae (mosses, not shown), rare shells and head shields of chydoridae and rare head capsules of larvae of chironomidae. the fossil assemblage and in particular the occurrence of abundant remains of telmatic plants show that the sediment was deposited in shallow water near the shore of a lake, probably just outside the reed belt. the presence of fruits of betula sect. albae indicates that the land was covered by birch forests. the sample of cladium mariscus fruits yielded an age of 10 980 ± 55 14c years bp (kia-21680). this is calibrated to 12.674–13.069 cal. ka bp, according to the intcal09 dataset, which corresponds to the youngest part of the allerød chronozone or the oldest part of the younger dryas. an age corresponding to the warm allerød chronozone was expected from the fossil assemblage, because cladium mariscus and scirpus lacustris are thermophilous plants. cladium mariscus was recorded from late-glacial deposits in south-eastern denmark by bennike & jensen (1995), but its presence was probably due to down-core contamination, and there are no secure records of it from late-glacial deposits in denmark (iversen 1954; jensen et al. 1997; bennike et al. 2004). its northern geographical limit during the allerød may thus have been located near the coring site. as the core was collected at a water depth of 39.5 m and the sample comes from a core depth of c. 5.8 m, the dated sample comes from a depth of c. 45.3 m below present sea level. the sediment is fine-grained and was probably deposited at a water depth of several metres, and we suggest that the shore level during deposition was around 40 m lower than at present. both before and after this lowstand episode, the relative shore level was around 20 m below the present sea level according to jensen et al. (1997). this implies that the shore-level fall towards the end of the allerød chronozone was of the same magnitude as the fall at the younger dryas – holocene transition, i.e. considerably more than 5–10 m as suggested by björck (1995). the new data allow us to modify the shore-level model proposed by bennike & jensen (1998) and extend it back in time. figure 4 shows a new model for relative shore-level changes in the arkona basin from the last deglaciation to the present. it is seen that transgressions were interrupted by sudden regressions. south of the arkona basin, late-glacial sediments reach elevations lower than 20 m (lampe 2005). however, the relationship between these sediments and the regional shore level of the southern baltic basin is uncertain. some of the sediments are glaciofluvial and were deposited above shore level, other late-glacial sediments may have been deposited in local basins, perhaps in part dammed by bodies of stagnant ice. at present, the most enigmatic stage in the history of the baltic basin is that of the early holocene ancylus lake; the shore-level curve for this stage has been drawn as a dashed line on fig. 4. it has been suggested that the ancylus regresfig. 4. tentative curve showing relative shore-level changes in the arkona basin during the late-glacial and the holocene. ka: kilo-annum (1000 years), bil: baltic ice lake, ys: yoldia sea, all: allerød, yd: younger dryas. modified from bennike & jensen (1998). table 1. macrofossils in core 258000 r: rare, c: common. 412–420 – – – – – – – – – – – c r – 420–430 – – – – – – – – – – – c r – 430–440 3 – – – – – – – – – – c r 4 440–449 2 – – – – – – – – – – c – 10 449–454 – – – – – – – – – – – c – 1 454–459 – – – – – – – – – – 5 c r 8 459–463 2 r – – – – – c – – 1 c r – 463–468 – – – – – – – c – – – c r – 468–473 1 r – – – – – c – 1 – c r – c. 580 3 – 50 15 150 20 1 c 200 1 – r r – be tu la se ct . a lba e sp . d ep th (c m ) pin us sy lve str is cl ad ium m ar isc us sc irp us la cu str is ca re x sp p. co m ar um p alu str e ci cu ta vi ro sa ra di ce lls m en ya nt he s t rif oli at a n ym ph ae a sp . pis cic ola g eo m et ra c hy do rid ae in de t. c hi ro no m id ae in de t. cr ist at ell a m uc ed o d ep th (m b .s. l.) 12 10 8 6 4 2 0 age (cal. ka bp) 0 10 20 30 40 littorina seaancylus lakeysbil focus of this paper holoceneydall 2020 sion was around 20 m (björck 1995), but more recently figures of 5 m and 10 m were also proposed (björck et al. 2008; rosentau et al. 2013). no erosional unconformity has been reported from the south-western baltic basin that formed during this regression (jensen et al. 1999). figure 5 shows a model of the palaeogeography of the baltic basin after the late allerød drainage. although the baltic basin was now at the same level as the sea, it probably remained a freshwater lake. the connection to the sea was narrow, and we suggest that the outflow of huge amounts of fresh water coming from rivers and from the melting ice sheet hindered seawater from entering the baltic basin. concluding remarks the rich occurrence of remains of reed plants at a depth of 45 m below sea level in the arkona basin provides firm evidence of a lowstand. a radiocarbon age shows that it occurred at the end of the allerød chronozone. we suggest that the shore level fell about 20 m, similar to the shore-level fall at the younger dryas – holocene boundary. acknowledgements the captain and crew of the former r/v alexander von humboldt from the institute for baltic sea research in warnemünde are thanked for their help during the marine cruise. this paper is dedicated to the memory of wolfram lemke, who invited us to take part in the cruise during which cores 258000-1 and 258000-2 were collected. references andrén, t., björck, s., andrén, e., conley, d., zillén, l. & anjar, j. 2011: the development of the baltic sea during the last 130 ka. in: harff, j. et al. (eds): the baltic sea basin, 75–97. berlin: springer verlag. bennike, o. & jensen, j.b. 1995: near shore baltic ice lake deposits in fakse bugt, southeast denmark. boreas 24, 185–195. bennike, o. & jensen, j.b. 1998: lateand postglacial shore level changes in the southwestern baltic sea. bulletin of the geological society of denmark 45, 27–38. bennike, o., jensen, j.b., lemke, w., kuijpers, a. & lomholt, s. 2004: lateand postglacial history of the great belt, denmark. boreas 33, 18–33. björck, s. 1995: a review of the history of the baltic sea, 13.0–8.0 ka bp. quaternary international 27, 19–40. björck, s., andrén, t. & jensen, j.b. 2008: an attempt to resolve the partly conflicting data and ideas on the ancylus–littorina transition. polish geological institute special papers 23, 21–26. iversen, j. 1954: the late-glacial flora of denmark and its relation to climate and soil. danmarks geologiske undersøgelse ii. række 80, 87–119. jensen, j.b., bennike, o., witkowski, a., lemke, w. & kuijpers, a. 1997: the baltic ice lake in the southwestern baltic: sequence-, chronoand biostratigraphy. boreas 26, 217–236. jensen, j.b., bennike, o., witkowski, a., lemke, w. & kuijpers, a. 1999: early holocene history of the southwestern baltic sea: the ancylus lake stage. boreas 29, 437–453. lampe, r. 2005: lateglacial and holocene water-level variations along the ne german baltic sea coast: review and new results. quaternary international 133–134, 121–136. larsen, c.s. 2004: sequence stratigraphy based on vibrocore description and shallow seismic data from the south-western baltic sea. danmarks og grønlands geologiske undersøgelse rapport 2004/53, 28 pp. lemke, w., endler, r., tauber, f., jensen, j.b. & bennike, o. 1998: late and postglacial sedimentation in the tromper wiek (western baltic). meyniana 50, 155–173. moros, m., lemke, w., kuijpers, a., endler, r., jensen, j.b., bennike, o. & gingele, f. 2002: regression and transgressions of the baltic basin reflected by a new high-resolution deglacial and postglacial lithostratigraphy for arkona basin sediments (western baltic sea). boreas 31, 151–162. rosentau, a. et al. 2013: stone age settlement and holocene shore displacement in the narva-luga klint bay area, eastern gulf of finland. boreas. http://dx.doi.org/10.1111/bor.12004 uścinowicz, s. 2006: a relative sea-level curve for the polish southern baltic sea. quaternary international 145–146, 86–105. wohlfarth, b., björck, s., funder, s., houmark-nielsen, m., ingólfsson, ó., lunkka, j.-p., mangerud, j., saarnisto, m. & vorren, t. 2008: quaternary of norden. episodes 31, 73–81. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: obe@geus.dk fig. 5. generalised palaeogeographical map of the baltic region after the drainage at the end of the allerød. modified from björck (1995) and wohlfahrt et al. (2008). b: billingen. 250 km ice sea lake land b 258000 http://dx.doi.org/10.1111/bor.12004 mailto:prj@geus.dk geological survey of denmark and greenland bulletin 1, 865-892 865 the jurassic of kuhn ø, north-east greenland per c. alsgaard,vince l. felt, henrik vosgerau and finn surlyk the middle–upper jurassic succession of kuhn ø, north-east greenland accumulated in a major half-graben and is an excellent analogue for the subsurface of the mid-norwegian shelf. on kuhn ø, peneplaned crystalline basement was incised by a drainage system during a major base-level lowstand, probably in late early or early middle jurassic times. it was filled with fluvial conglomerates of the newly defined middle jurassic bastians dal formation during subsequent base-level rise. as sea level continued to rise, precursor-peat of the coals of the muslingebjerg formation formed in swamps which covered the conglomerates and filled the remaining space of the incised valley system. the valley and interfluve areas were flooded in late bathonian – callovian times and tidally-dominated, shallow marine sandstones of the pelion formation were deposited on top of the valley fill and over the adjacent basement peneplain. these sandstones are overlain by the newly defined shallow marine oxfordian payer dal formation which is subdivided into a lower unit and an upper unit, separated by a major drowning surface. the payer dal formation sands were flooded in the late jurassic and organic-rich, offshore mudstones of the bernbjerg formation were deposited. the jurassic succession of kuhn ø can thus be subdivided into large-scale sedimentary units separated by major drowning surfaces. they are of regional extent, and in combination with biostratigraphic and 87sr/86sr isotope data they allow the correlation of the sedimentary units on kuhn ø with more offshore deposits to the south in wollaston forland and more landwards successions to the north in hochstetter forland. petrographically, the trough cross-bedded sandstones of the pelion formation and the lower unit of the payer dal formation include both calcite-cemented and poorly cemented quartz sandstones. the calcite cement was derived from dissolution of abundant calcareous fossils and forms concretionary horizons. the upper unit of the payer dal formation mainly consists of weaklycemented quartz sandstones with porosities around 30%. the sandstones of the pelion and payer dal formations on kuhn ø are petrographically very similar to jurassic sandstones from the midnorwegian shelf and the barents sea with regard to original mineralogical composition, sorting and grain size. the bernbjerg formation mudstones are comparable to the upper jurassic source rock of the mid-norwegian shelf and the barents sea, but have lower hydrogen index (hi) values due to terrigenous input in a relatively proximal setting. coals of the muslingebjerg formation have significant source rock potential with measured hi values up to 700, kerogen types ii–iii and total organic carbon (toc) values above 50%. keywords: kuhn ø, north-east greenland, middle–upper jurassic, lithostratigraphy, sedimentology, petrography, source rocks p.c.a.* & v.l.f.‡, amoco norway oil company. present addresses: *norsk hydro, n-0246 oslo, norway. ‡ bp amoco – egypt, 14, road 252, digla, maadi cairo, egypt; p.o. box 2409. e-mail: per.chr.alsgaard@hydro.com h.v., geological survey of denmark and greenland, geocenter copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. f.s., geological institute, university of copenhagen, geocenter copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. geological survey of denmark and greenland bulletin 1, 865–892 (2003) © geus, 2003 the proximity of north-east greenland to the midnorwegian shelf during mesozoic times makes it an obvious candidate in the search for exposed analogues of reservoir facies, source rocks and structural framework of the potential hydrocarbon systems. kuhn ø, in particular, may be analogous to areas on the midnorwegian shelf that are characterised by a relatively thin jurassic succession between basement and the ‘base-cretaceous unconformity’ that accumulated adjacent to a subaerially exposed landmass. the ‘wollgan project’ was a co-operative research project involving statoil, saga, amoco, the norwegian petroleum directorate (npd) and the geological institute of the university of copenhagen. the three weeks of field work in 1994 focused on the jurassic early rift and rift-climax successions of kuhn ø and wollaston forland, and the cretaceous post-rift deposits of wollaston forland (figs 1, 2). within this framework, geologists from amoco and the university of copenhagen studied the jurassic of kuhn ø, forming the basis for this paper. the field work took place in central and south kuhn ø (figs 3–5). the geology was mapped onto non-orthographic vertical aerial photographs enlarged to a scale of 1:25 000. the middle–late jurassic structural development of tilted fault blocks in the kuhn ø – wollaston forland region was first unravelled by vischer (1943). maync (1947) documented the jurassic and lower cretaceous sections of kuhn ø and identified many good outcrops. the geological map by koch & haller (1971) gives a general structural and stratigraphic overview, and a lithostratigraphic scheme was erected by surlyk (1977). 866 milne land traill ø geographical society ø clavering ø kuhn ø store koldewey 100 km 22°w 18°w 26°w 22°w 72°n 74°n 76°n normal fault reverse fault fault, indeterminate type hochstetter forland wollaston forland hold with hope jameson land k tt c greenland fig. 1. map of the east greenland basin complex showing fault zones active during mesozoic times. c, cardiocerasdal; k, kulhus, søndre muslingebjerg; tt, th. thomsen land. groupseries formation member bernbjerg niesen rigi laugeites ravine payer dal pelion muslingebjerg supergroup ja m es on l an d wollaston forland hall bredning vardekløft middle jurassic upper jurassic lower cretaceous jakobsstigen bastians dal lindemans bugt palnatokes bjerg rødryggen albrechts bugt falske bugt young sund ugpik ravine spath plateau fig. 2. stratigraphic scheme for the jurassic–cretaceous of hold with hope, wollaston forland, kuhn ø and hochstetter forland; slightly modified from surlyk (2003, this volume, fig. 5). note that the jakobsstigen formation is restricted to wollaston forland and th. thomsen land and is not recognised on kuhn ø. it is laterally equivalent to the lower payer dal formation on kuhn ø (see fig. 23). 867 fligely fjord bastians d al ? ? ? kap hamburg kap maurer ba sti an b ug t kingofjeld schwarze wand bernbjerg pa ye r d al cross-se ction 5 km n 043 c-58 c-57 157 c-87, c-88 baselbjerg quaternary alluvium basaltic lavas/intrusions wollaston forland group and younger strata bernbjerg formation palaeogene upper jurassic – lower cretaceous upper jurassic payer dal fm (upper unit) pelion fm and payer dal fm (lower unit) (boundary indicated where differentiation possible) bastians dal and muslingebjerg fms caledonian crystalline basement sample locality measured section stratigraphic boundaries observed stratigraphic boundaries inferred normal fault, tick on downthrow side inferred fault line peaks strike and dip middle–upper jurassic basement fig. 3. geological map of kuhn ø, east greenland. the cross-section indicated is shown in fig. 7. the two areas outlined show the location of the detailed maps in figs 4, 5. based on koch & haller (1971) and own data. surlyk & clemmensen (1983) interpreted the sedimentary succession in terms of a series of backstepping units reflecting the combined effects of progressive rifting and eustatic sea-level rise. surlyk (1991) interpreted the middle–upper jurassic succession within a low order sequence stratigraphic framework. this paper presents a revised lithostratigraphic scheme for the jurassic of kuhn ø and two new formations are defined, the fluvial bastians dal formation at the base of the succession and the shallow marine payer dal formation (fig. 2). the jurassic sediments are described according to their facies, diagenesis and geochemistry, and an interpretation of the depositional environments is presented. thickness and facies changes from the crest of a tilted fault block in the east (schwarze wand) to a relatively deeper structural setting on the hangingwall in the west (west payer dal) is illustrated by an east–west stratigraphic profile through southern kuhn ø, parallel to structural dip. finally, the jurassic succession of kuhn ø is compared to more offshore contemporaneous deposits to the south in wollaston forland and a more landwards succession to the north in hochstetter forland described by clemmensen & surlyk (1976), surlyk (1977, 1978a), surlyk & clemmensen (1983), bojesen-koefoed et al. (1996), petersen et al. (1998) and vosgerau et al. (2000). 868 n 2 km c-10 c-8 c-6 c-13 c-20 c-21c-22 c-23, c-24 c-2 c-59 c-3 c-43 c-51 c-54 c-49 037 c-47 c-46c-11 c-38 c-42 c-41 c-39 c-40 c-15 c-16 c-17 c-17a c-60, c-62 ? ? ? a b4 b3 c d3 e d4 d2 d1 b2 b1 c-19 11 11 12 kap hamburg ugpik ravine kingofjeld schwarze wand bernbjerg pa ye r d al fig. 4. geological map of the payer dal area, south kuhn ø (see fig. 3). based on koch & haller (1971) and own data. for legend, see fig. 3. geological setting the jurassic succession of kuhn ø was deposited in the northern part of the wollaston forland basin of north-east greenland, which was situated on the western margin of the jurassic rift complex between greenland and norway. deposition took place in a rift-controlled embayment which was open to marine circulation towards the south (fig. 6). regional sediment transport was axial from north to south, down a low-gradient basin floor sloping to the south (surlyk 1977, 1990, 1991; surlyk & clemmensen 1983). the island of kuhn ø is positioned on the hangingwall slope of a tilted fault block, within which middle jurassic strata rest unconformably on crystalline basement (fig. 7). the jurassic succession thickens westwards into the half-graben and onlaps eastwards onto caledonian basement. the middle–upper jurassic succession of kuhn ø comprises, in ascending order, the bastians dal, muslingebjerg, pelion, payer dal and 869 2 km c-74 c-79, c-80a c-66 c-85 c-67, c-68, c-69 c-77, c-78 fligely fjord bastians dal 6 5–17 14–30 n baselbjerg fig. 5. geological map of the bastians dal area, central kuhn ø (see fig. 3). based on koch & haller (1971) and own data. for legend, see fig. 3. bernbjerg formations which form a series of backstepping early rift units separated by major drowning surfaces (figs 2, 8; surlyk 1977, 1991; surlyk & clemmensen 1983). it is overlain by conglomerate-dominated rift-climax deposits of the wollaston forland group of latest jurassic – earliest cretaceous age (surlyk 1978b). methods the jurassic sediments were studied in southern kuhn ø around the valley of payer dal and in central kuhn ø around the valley of bastians dal (figs 3–5). the exposed section extends from the basal unconformity overlying crystalline basement, through the bastians dal, muslingebjerg, pelion, payer dal and bernbjerg formations (figs 8–11). the outcrops are partially covered by scree slopes, such that the sections are composite and combined from smaller segments. a total of 93 samples were collected on kuhn ø for petrographic, geochemical and biostratigraphic analysis. to avoid the effects of surface weathering, geochemical samples were normally collected at depths of 30 cm into the permafrost, or almost one metre from the surface. a total of 33 samples were evaluated biostratigraphically, ten for calcareous nannoplankton, nineteen for foraminifers and four for ammonites (table 1). dinocysts are absent in the non-marine bastians dal formation; the marine pelion formation and the lower payer dal formation show higher dinocyst diversities than the upper payer dal formation. the samples analysed for micropalaeontology were barren of nannoplankton and there were only a few poorly preserved foraminifers with long stratigraphic age ranges. ammonites were only found in the bernbjerg formation. previous biostratigraphic data (sykes & surlyk 1976; surlyk 1977) were utilised in dating and correlating the jurassic succession on kuhn ø with adjacent areas within the wollaston forland basin. in addition, strontium isotope analysis has been performed on a few belemnites from the jurassic succession in payer dal and from a more offshore succession in the cardiocerasdal 870 22°w 18°w 76°n 75°n 25 km hochstetter forland kuhn ø wollaston forland paralic estuary with tidal shoals offshore, marine th. thomsen land fligely fjord bernbjerg payer dal a b c d ekingofjeld schwarze wand d ep th /a lti tu de ( km ) sw (250°) ne bernbjerg fm, wollaston forland group and lower cretaceous payer dal fm (upper unit) pelion fm and payer dal fm (lower unit) caledonian crystalline basement 1 0 -1 0 5 10 15 20 km fig. 6. late bathonian – callovian palaeogeography of the wollaston forland basin (modified from surlyk & clemmensen 1983). fig. 7. cross-section of south kuhn ø (for location, see fig. 3). the profile intersects measured sections a and e, while the other sections are projected onto the line of section. 871 bernbjerg fm (?upper oxfordian – kimmeridgian) hall bredning gp vardekløft gp upper jurassic middle jurassic series caledonian crystalline basement group upper unit (upper oxfordian) pa ye r d al f or m at io n lower unit (lower–middle oxfordian) pelion fm (upper bathonian – upper callovian) muslingebjerg fm bastians dal fm formation c la y si lt sa nd g ra ve l coal mudstone mudstone/sandstone heterolith sandstone ( calcareous cement) conglomerate/pebbly sandstone planar cross-bedding composite planar cross-bedding trough cross-bedding hummocky cross-stratification parallel lamination/stratification wavy bedding sand lenses (cross-laminated) shell-rich wave ripples (cross-lamination) current ripples (cross-lamination) siltstone clast fossil wood belemnite bivalve oysters horizontal burrow vertical burrow calcite concretion oyster bed 50 m fig. 8. generalised stratigraphic column of the jurassic early syn-rift succession of kuhn ø. the accompanying legend also applies to figs 11, 17. 872 ss aa pp aayy er d al er d al d yk d yk ee b er n b je rg b er n b je rg b 3, b 3, b 4 b 4 b 1, b 1, b 2 b 2 b er n b je rg f m b er n b je rg f m tto p p o p p aayy er d al f m er d al f m tto p lo o p lo ww er u n it er u n it pp aayy er d al f m er d al f m nn sa a p ay er d al d yk e b er n b je rg b 3, b 4 b 1, b 2 b er n b je rg f m to p p ay er d al f m to p lo w er u n it p ay er d al f m n li th os tr at ig ra ph ic b ou nd ar y n or m al fa ul t m ea su re d se ct io n ssbb s ch w ar z s ch w ar ze e ww an d an d tto p lo o p lo ww er u n it er u n it pp aayy er d al f m ? er d al f m ? b as em en t b as em en t pp aayy er d al er d al tto p lo op lo ww er u ni t er u ni t pp aayy er d al f m er d al f m tto p p op p el io n f m el io n f m cc tto p p o p p aayy er d al f m er d al f m b er n b je rg f m b er n b je rg f m pp al ae o al ae o gge n e b as al ts en e b as al ts tto p ba se m en t op b as em en t tto p p op p el io n fm ? el io n fm ? k in g k in go fje ld o fje ld nn sb s ch w ar ze w an d to p lo w er u n it p ay er d al f m ? b as em en t p ay er d al to p lo w er u ni t p ay er d al f m to p pe lio n f m c to p p ay er d al f m b er n b je rg f m p al ae og en e b as al ts to p ba se m en t to p pe lio n fm ? k in go fje ld n ? fi g. 9 . a : w es t si d e o f p ay er d al , so u th k u h n ø , sh o w in g th e p o si tio n s o f se ct io n s a an d b 1 –4 . t h e re fe re n ce s ec tio n o f th e p ay er d al f o rm at io n i llu st ra te d b y su rl yk ( 19 77 , fig . 4) w as m ea su re d i n t h e so u th er n m o st g u lly ( u gp ik r av in e, a rr o w ). t h e u p p er u n it o f th e p ay er d al f o rm at io n i s c. 6 0 m t h ic k. b : e as t si d e o f p ay er d al . t h e ri gh th an d p ar t o f th e p an o ra m a ill u st ra te s th e b es t ex p o su re s o f th e p el io n a n d p ay er d al f o rm at io n s o n s o u th k u h n ø a n d i n cl u d es t h e ty p e se ct io n o f th e p ay er d al f o rm at io n ( se ct io n c , se e al so f ig s 16 , 17 ). t h e lo w er u n it o f th e p ay er d al f o rm at io n i s c. 9 0 m t h ic k. 873 fi g. 1 0. a : sc h w ar ze w an d o n s o u th k u h n ø s h o w in g, t o t h e ri gh t (n e ), t h e cr es t o f a ju ra ss ic r o ta te d f au lt b lo ck . n o te t h e n o rt h -e as tw ar d s o n la p a n d t h in n in g o f th e se d im en ta ry p ac ke t b et w ee n b as em en t an d t h e to p o f th e lo w er p ay er d al f o rm at io n ; th is i n te rv al i s c. 4 0 m t h ic k at t h e n o rt h -e as t en d o f th e ill u st ra te d s ec tio n ( se e fi g. 1 0b ). t h e h ill i s ca p p ed b y p al ae o ge n e b as al ts . fo r le ge n d , se e fi g. 9 a . b : c lo se -u p o f th e n o rt h -e as te rn p o rt io n o f th e se ct io n i n f ig . 10 a ; a co m p o si te s ec tio n ( se ct io n e ) w as m ea su re d a lo n g th e d o tt ed lin es . c : p an o ra m a o f th e p en ep la n ed t o p o f th e cr ys ta lli n e b as em en t o ve rl ai n b y th e b as tia n s d al ( ?) , p el io n a n d p ay er d al f o rm at io n s o n t h e ea st s id e o f k in go fje ld ; th e ju ra ss ic se ct io n i s ca p p ed b y p al ae o ge n e b as al ts . fo r le ge n d , se e fi g. 9 a . s w s w a b tto p ba se m en t o p ba se m en t tto p lo o p lo ww er u ni t er u ni t pp aayy er d al f m er d al f m to p lo w er u ni t p ay er d al f m tto p lo o p lo ww er u n it , er u n it , p p aayy er d al f m er d al f m to p lo w er u n it , p ay er d al f m to p b as em en t p al ae og en e b as al ts n e n e to p ba se m en t s ec to p b as em en t p al ae o ge n e b as al ts tto p lo o p lo ww er u n it er u n it pp aayy er d al f m er d al f m to p lo w er u n it p ay er d al f m n w area, south-western wollaston forland (table 2; m. engkilde, personal communication 1997). the 87sr/86sr isotope values were used for correlating the jurassic succession between the two areas and with the strontium isotope curves constructed for the jurassic period for the united kingdom (jones et al. 1994) and east greenland (m. engkilde, personal communication 1997). 874 formation age sample* diagnostic flora/fauna – mainly dinocysts, together with pollen (p), ammonites (a), foraminifers (f) early–middle barremian c-57 pseudoceratium anaphrissum, muderongia australis, pseudoceratium pelliferum, muderongia staurota, muderongia testudinaria early–middle barremian c-58 p. anaphrissum, batioladinium longicornutum bernbjerg late callovian – oxfordian c-22 scriniodinium crystallinum bernbjerg middle–late jurassic c-23 cerebropollenites mesozoicus (p), haplophragmoides sp. (f) bernbjerg kimmeridgian c-24 aulacostephanus eudoxus (a) bernbjerg bathonian–kimmeridgian c-41, c-42, c-79 gonyaulacysta jurassica bernbjerg kimmeridgian–volgian c-74 rhynchodiniopsis cladophora, g. jurassica, oligosphaeridium pulcherrimum, ammobaculites sp. (f), haplophragmoides sp. (f) bernbjerg oxfordian–kimmeridgian c-77, c-78 r. cf. r. cladophora bernbjerg kimmeridgian c-80a aulacostephanus mutabilis (a) bernbjerg kimmeridgian–volgian c-87 tubotuberella apatela, ?atopodinium sp., sirmiodinium grossii, hystrichodinium cf. amphiacanthum. bernbjerg volgian c-88 oligosphaeridium patulum, scriniodinium sp. a, cribroperidinium sp. of the perforans/cauda group bernbjerg late callovian – oxfordian, 037, sect. d4 escarisphaeridum sp. hyalina, sentusidinium pelionense, probably late oxfordian sirmiodinium grossii, g. jurassica bernbjerg kimmeridgian 157 oligosphaeridium patulum, ellipsoidictyum cinctum, mendicodinium groenlandicum, c. mesozoicus (p) bernbjerg kimmeridgian a4 sect. d4 amoeboceras sp. baylei (a) bernbjerg kimmeridgian 65 m below c-87 amoeboceras subkitchini spath (a) payer dal (upper unit) middle–late jurassic c-15, c-49 c. mesozoicus (p) pelion/payer dal (lower unit) oxfordian c-3p g. jurassica, m. groenlandicum, rigaudella aemula pelion/payer dal (lower unit) middle–late jurassic c-10, c-11, c-21 c. mesozoicus (p) pelion/payer dal (lower unit) callovian–oxfordian c-13 sentusidinium rioultii, rhynchodiniopsis cf. cladophora payer dal (lower unit) late callovian – c-38 liesbergia scarburghensis, g. jurassica, ?surculosphaeridium middle oxfordian vestitum, r. cladophora pelion/payer dal (lower unit) late callovian – c-46 l. scarburghensis middle oxfordian pelion/payer dal (lower unit) callovian c-47 pareodinia prolongata pelion/payer dal (lower unit) late bathonian – c-54 batiacasphaera dictydia, sirmiodinium grossii, g. jurassica, late callovian g. pectinigera, escharisphaeridia sp., molluscs, brachiopods pelion/payer dal (lower unit) ?callovian c-59 chytroeisphaeridia chytroeides pelion/payer dal (lower unit) middle–late jurassic 043 c. mesozoicus (p), ?rhynchodiniopsis cladophora payer dal ?callovian c-66 c. chytroeides pelion middle–late jurassic c-85 c. mesozoicus (p) bastians dal middle–late jurassic c-19, c-68, c-69 c. mesozoicus (p) * location of samples shown on figures 3–5, 11, 17. table 1. biostratigraphic data, jurassic – lower cretaceous of kuhn ø stratigraphy and sedimentology the jurassic succession on kuhn ø is subdivided into five formations, the bastians dal (new), muslingebjerg, pelion (redefined), payer dal (new) and bernbjerg formations (fig. 2). the first four formations occur within the vardekløft group whereas the bernbjerg formation is referred to the hall bredning group (fig. 2). the new and redefined formations are described formally below, together with a brief account of the characteristic features of the muslingebjerg and bernbjerg formations on kuhn ø. it should be noted that the revised lithostratigraphic scheme presented in figure 2 and in surlyk (2003, this volume, fig. 5) must be considered as provisional, pending publication. in this paper, this relates especially to the muslingebjerg and pelion formations which are described here from kuhn ø, as these units are elevated in status from member to formation in the revised scheme. bastians dal formation new formation history. a unit of cobble conglomerates, 5 m thick, at the base of the pelion formation and directly overlying crystalline basement was described from one locality on the mountain of kingofjeld, eastern payer dal by maync (1947, p. 15). this unit is included in the new bastians dal formation which is erected here for a succession of mainly conglomerates and pebbly sandstones forming the basal unit of the jurassic succession on kuhn ø. the type area of the formation was mapped by vischer (1943, p. 42–43) but rocks belonging to the new bastians dal formation were not recognised, probably due to deep snow cover (a. vischer in: koch 1955, p. 562–564). name. after the nw–se-trending valley of bastians dal, in western central kuhn ø (figs 3, 5; vischer 1943, plate 1). 875 c-85 locally up to c. 150 m thick conglomerate c-68 c-69 c-67 30 m pelion formation muslingebjerg formation bastians dal formation basement 20 0 c la y si lt v f f m c v c p c gravel b sand fig. 11. stratigraphic log of the bastians dal succession, including the type section of the bastians dal formation. the uppermost part of the pelion formation is not shown. for legend, see fig. 8; c-67, sample number. 876 type locality and type section. the east side of bastians dal, where the type section is located (figs 5, 11). thickness. maximum thickness of 100–150 m in bastians dal, estimated from photographs and measured sections; the formation thins rapidly towards the south within the bastians dal area and is only 1–5 m thick on southern kuhn ø (maync 1947). lithology. the bastians dal formation consists of quartz pebble conglomerates and pebbly sandstones. the formation has a distinctive grey weathering colour, due to a high content of carbonaceous material and coal lenses, and contrasts with the light brown weathering colour of the pelion formation above (fig.12). boundaries. the formation rests directly on crystalline basement rocks on kuhn ø. the upper boundary is placed where the conglomerates and pebbly sandstones of the bastians dal formation are overlain by coals of the muslingebjerg formation or, where the muslingebjerg formation is not developed, by fineto medium-grained sandstones of the pelion formation. distribution. the formation overlies crystalline basement in central and southern kuhn ø. in bastians dal, the fig. 12. grey fluvial conglomerates of the bastians dal formation overlain by light brown shallow marine sandstones of the pelion formation (c. 20 m thick pelion formation section crops out along the ridge, centre foreground). bastians dal, central kuhn ø, viewed towards the north with th. thomsen land in the background. arrow indicates location of exposed conglomerates illustrated in fig. 13. fig. 13. stacked fining-upwards conglomeratic units with cut and fill structures, typical of the bastians dal formation. hammer, 30 cm long, for scale. for location, see fig. 12. 877 formation seems to fill a valley incised into the crystalline basement whereas in the payer dal area, it is thin and locally absent. geological age. the bastians dal formation is of general middle jurassic age based on the presence of cerebropollenites mesozoicus pollen within the formation and the occurrence of dinocysts within the overlying pelion formation which indicate a late bathonian – late callovian age (table 1). facies. in bastians dal, the formation is characterised by stacked fining-upwards units, 0.1–2 m thick, consisting of coarse-grained pebble and cobble conglomerates with quartzite clasts overlying a basal erosional surface followed by trough cross-bedded or parallel-laminated quartz sandstones, commonly rich in mica (fig. 13). cobbles are rounded whereas sand grains are angular to subrounded (fig. 14a). the poor sorting and angularity of the quartz grains suggest that the sediments are immature and close to their source of origin, probably the crystalline basement rocks to the east and north. microfractures are present in all quartz grains, supporting a metamorphic basement origin. palaeocurrent directions measured from the trough cross-bedded sandstones are towards the south-west (average 240°). palaeoenvironment. the carbonaceous, immature conglomerates and pebbly sandstones of the bastians dal formation are interpreted to have been deposited by fig. 14. thin-section photographs in plane-polarised light. a: sandstone from the uppermost bastians dal formation illustrating the poor sorting, angularity and poorly cemented character of the deposits. sample c-67, bastians dal. b: weakly cemented sandstone from the payer dal formation (lower unit) with 19% helium porosity; sample c-60, west kingofjeld (fig. 17). c: cemented sandstone (cf. fig. 14b) from the payer dal formation (lower unit); sample c-62, west kingofjeld (fig. 17). a c b 1mm 1mm 1mm 878 south-westwards flowing braided rivers. the finingupwards units reflect decreasing energy conditions during deposition and were probably formed by fluvial channel erosion followed by the migration of channel bars towards the south-west within the channels (cant & walker 1976; miall 1977). the formation appears to be mainly restricted to the bastians dal area (fig. 3). the overall marked peneplaned nature of the basement surface and the pronounced lateral thinning of the formation suggest that it is restricted to a valley system incised into the basement surface. it is not possible to define the orientation of the axis of the valley or the northwards extent of the formation due to quaternary cover. incision is interpreted to have resulted from relative base-level lowering, probably in association with late early or early middle jurassic regional uplift enhanced by the onset of half-graben block rotation in middle jurassic time. incised drainage systems possibly followed zones of weakness in the crystalline basement such as fault zones or less resistant metamorphic facies. during a subsequent rise in base level, a south-west flowing fluvial system filled the valley with sandstones and conglomerates. on southern kuhn ø, the bastians dal formation was apparently deposited on the peneplain marginal to the incised valley system inferred for central kuhn ø. muslingebjerg formation facies. coal beds overlying the bastians dal formation are referred to the coal-bearing muslingebjerg formation. this lithostratigraphic unit was given member status by 5 m 20 m coal coal pelion fm (fine-grained sandstone) very coarsegrained sandstone river level fig. 15. exposure of the coal-dominated muslingebjerg formation, overlain by fine-grained, well-sorted marine sandstones of the pelion formation; bastians dal, central kuhn ø. as illustrated in the sketch (no vertical exaggeration), the coals interdigitate laterally with very coarse-grained fluvial sandstones (arrow in photograph), comparable to those of the underlying bastians dal formation. the sandstone is partially covered by coal dust and is thus difficult to see on the photograph. surlyk (1977) but is upgraded to formation in the revised scheme of surlyk (2003, this volume, fig. 5). at the type locality at kulhus in southern hochstetter forland, the formation is at least 20 m thick and consists of coals interbedded with lagoonal siltstones and sandstones and subordinate shoreface sandstones (clemmensen & surlyk 1976; petersen et al. 1998). in bastians dal, the muslingebjerg formation is up to 11 m thick and consists of coals interbedded locally with fluvial deposits comparable to those of the underlying bastians dal formation (figs 11, 15). in payer dal, a coal bed up to 0.15 m thick overlies weathered crystalline basement or thin fluvial deposits of the bastians dal formation. the coal beds are overlain by fineto medium-grained sandstones of the pelion formation on kuhn ø and the payer dal formation on hochstetter forland. the muslingebjerg formation on kuhn ø is of early middle jurassic age based on the occurrence of middle–upper jurassic pollen in the underlying fluvial sediments of the bastians dal formation and upper bathonian – upper callovian dinocysts in the overlying pelion formation (table 1). dinocysts immediately above the coals in the type section at kulhus, hochstetter forland, are indicative of the upper callovian p. athleta chronozone (piasecki & stemmerik in press). palaeoenvironment. the coal beds of the muslingebjerg formation represent lagoonal swamps that formed on a coastal plain during base-level rise. the alternation of coal beds and shoreface deposits at kulhus demonstrates that the peat-forming mires were repeatedly terminated by marine transgressions (clemmensen & surlyk 1976; petersen et al. 1998). interfingering of coals with fluvial deposits in bastians dal on kuhn ø suggests that the coal beds were also confined within the incised valley system within which the fluvial bastians dal formation was deposited. the thin coal beds that directly overlie or occur close to basement in payer dal were possibly deposited immediately adjacent to the eastern low-angle margin of the incised valley. as flooding continued, the incised valley and interfluves were drowned and the coal-bearing sediments were abruptly overlain 879 se nw ttop loop lowwer uniter unit ppaayyer dal fmer dal fmttop loop lowwerer unitunit ppaayyer daler dal fmfm ttop pop pelion fmelion fm ffaultault top lower unit payer dal fmtop lower unit payer dal fm top pelion fm top payer dal fm (projected) fault fig. 16. type section (section c) of the payer dal formation, west kingofjeld (see figs 4, 9b, 17); the line of section is indicated by the dotted lines. the boundary between the payer dal formation and the overlying bernbjerg formation is exposed about one kilometre east (left) of this locality (see fig. 9b). the lower unit of the payer dal formation is c. 90 m thick. for legend, see fig. 9a. by shallow marine sands of the pelion or payer dal formations. the muslingebjerg formation thus marks the transition from fluvial sedimentation of the bastians dal formation to fully marine sedimentation of the pelion and payer dal formations. geochemistry. coal samples from payer dal have high hydrogen indices (hi) in the range 430–695 mg hc/g toc (mg hydrocarbons/g total organic carbon content) (fig. 4; samples c-6, -8, -19, -20). values of tmax are in the range 429–432°c, and toc values are 51–59%. a coal sample from central kuhn ø (fig. 5, c-68) gave a hi value of only 170 mg hc/g toc. the coals are relatively ‘liptinitic’, corresponding to kerogen types ii–iii. the organic matter is dominantly derived from land plants, and has a good potential for generation of probably waxy oil. the coals of hochstetter forland are also potential oil-prone source rocks (bojesen-koefoed et al. 1996; petersen et al. 1998). pelion formation redefined history. the shallow marine sandstones on kuhn ø that overlie crystalline basement or in some areas succeed the bastians dal or muslingebjerg formations were previously all included in the pelion member (surlyk 1977); this lithostratigraphic unit is promoted to the rank of formation in the revised lithostratigraphic scheme (surlyk 2003, this volume, fig. 5). the sandstones are here subdivided into the pelion formation (below) and the new payer dal formation (above) which are separated by a major drowning surface expressed as a sharp boundary between mainly medium-grained sandstones and the overlying siltstones, heteroliths and fine-grained sandstones. the description presented here only covers the occurrence on kuhn ø. thickness. about 120 m in payer dal. lithology. the pelion formation consists of yellowish, moderately to well-sorted, fineto coarse-grained quartz sandstone. 880 c-60 c-62 c-38 160 150 scree slope scree slope scree slope scree slope scree slope 130 120 110 100 80 70 60 20 m pelion fm payer dal fm (lower unit) pd(u) c la y si lt f m c g r sand fig. 17. section c through the pelion – lower payer dal formations, west kingofjeld, including the type section of the payer dal formation (figs 9b, 16). for legend, see fig. 8. pd (u), payer dal formation, upper unit; c-38, sample number. 881 n a b c d e 1. 7k m 2.8km 2.7km 5. 4k m kuhn ø datum top basement top basement ? e d c b a 100 m n a b c d e 1. 7 km 2.8 km 2.7 km 5. 4 km kuhn ø bernbjerg formation payer dal formation (upper unit) payer dal formation (lower unit) pelion formation ?bastians dal formation ? ? ? m ud f m c v c m ud f m c v c m ud f m c v c m ud f m c v c m ud f m c v c sand sand sand sand sand mudstone sandstone fig. 18. west–east stratigraphic profile of the jurassic in southern kuhn ø. the datum is defined by the boundary between the lower and upper units of the payer dal formation. according to this correlation, the succession beneath the datum (pelion formation and lower payer dal formation) thins markedly eastwards. section c (type section of the payer dal formation) is shown in detail in fig. 17. section d is composite, constructed from segments d1–4 (see fig. 4). boundaries. overlies crystalline basement or pebbly fluvial sandstones of the bastians dal formation, and the coal-bearing muslingebjerg formation in southern and central kuhn ø. the upper boundary is sharp and separates pelion formation sandstones from siltstones and heteroliths of the overlying payer dal formation. it can be correlated with a similar surface at the top of the pelion formation in cardiocerasdal, wollaston forland, and can possibly be traced as far south as jameson land. the pelion formation and the lower unit of the payer dal formation cannot be clearly separated in bastians dal, central kuhn ø. distribution. the pelion formation is persistent throughout the jurassic outcrop on kuhn ø and regionally is recognised from store koldewey in the north (piasecki et al. in press) to jameson land in the south (fig. 1). geological age. the formation is poorly dated on kuhn ø due to a lack of ammonites. however, a late bathonian ammonite fauna in the basal part of the formation was reported from nearby wollaston forland by maync (1947). dinocyst assemblages suggest a late bathonian – late callovian age on kuhn ø (table 1). strontium isotope analysis (87sr/86sr) of a belemnite from the uppermost part of the formation suggests a late or possibly middle callovian age, by comparison with the strontium isotope curve of jones et al. (1994; m. engkilde, personal communication 1997). facies. the formation consists mainly of moderately well-sorted quartz sandstones with subangular to subrounded grains. ammonites are very rare, whereas belemnites and bivalves are abundant at certain levels. faunal assemblages are generally of low diversity and high density probably indicating a high-stress marine environment with fluctuating salinities (surlyk 1977). in the payer dal area, the pelion formation is best exposed in section c, west kingofjeld, where it is about 120 m thick (figs 4, 9b, 16–18). the formation is dominated by evenly laminated, swaley cross-stratified, and intensely bioturbated, fineto medium-grained sandstones alternating with wave rippled or trough cross-bedded, mediumto coarse-grained sandstones. structureless pebbly sandstone lags with erosional lower boundaries, sometimes rich in bivalve shells, are present locally. a calcite-cemented, bench-forming sandstone (c. 15 m thick), which seems to be of regional extent on kuhn ø, forms the top of the formation and is capped by the marked flooding surface (see above). in the more proximal setting in bastians dal, the pelion and payer dal formations are partially covered by scree and it has not been possible to recognise the surface separating the two formations. shallow marine sandstones probably occur both below and above the drowning surface and good outcrops are required in order to differentiate the two formations. the lower 45 m of the pelion formation in bastians dal consist of well-sorted, very fineto fine-grained sandstones with numerous thin carbonaceous laminae. layers of pebbles and poorly sorted beds ranging in grain size from fine sand to cobbles occur locally. only few sedimentary structures and bedding planes can be observed due to weathering and scree cover, with the exception of small-scale trough cross-beds and, towards the top, localised herringbone cross-bedding. this basal unit is probably overlain by a sandstone succession, estimated to be c. 35 m thick, which is exposed further to the west. correlation is uncertain, however, due to the intervention of an inferred nnw–ssetrending fault. the succession consists of four prominent 882 fig. 19. calcite concretions in the pelion formation or the lower payer dal formation. bastians dal, central kuhn ø. exposed sections, each c. 3–8 m thick, that comprise light brown, concretionary sandstones overlain by light grey, calcite-cemented sandstones; these benches are separated by poorly exposed, scree-covered intervals of similar thickness. the concretionary sandstones display characteristic ‘cannon ball’ concretions which litter the slopes below (fig. 19). the concretions, which range from a few centimetres to 30 cm in diameter (average 10 cm), commonly obscure sedimentary structures. the calcite-cemented sandstone intervals show trough crossbedding and are typically capped by structureless or parallel-laminated beds. the upper part of the formation is poorly exposed in this area. palaeoenvironment. onset of deposition of the pelion formation records regional marine flooding of the crystalline basement peneplain or the sediments of the bastians dal and muslingebjerg formations. deposition took place in the lower to upper shoreface based on the dominance of evenly laminated, swaley cross-stratified, and intensely bioturbated fineto medium-grained sandstones alternating with wave rippled or trough cross-bedded, mediumto coarse-grained sandstones. the pebbly sandstone lags with erosional lower boundaries which occur locally in the formation are interpreted as transgressive lags, formed by wave winnowing of underlying upper shoreface and foreshore deposits. a detailed account of the sedimentology and sequence stratigraphy of the pelion formation further south in jameson land is given by engkilde & surlyk (2003, this volume). payer dal formation new formation history. the rocks of this new formation were first recognised by maync (1947) who included them in the upper part of his yellow series. they form the upper part of the pelion member of surlyk (1977). name. after the valley of payer dal in southern kuhn ø (figs 4, 9). type locality and type section. the eastern side of payer dal; the type section (section c) is located on the west flank of kingofjeld (figs 4, 9b, 16, 17). reference section. ugpik ravine, west payer dal (fig. 9a; surlyk 1977, fig. 4). thickness. in payer dal, the formation is about 150 m thick. the thickness of the formation in bastians dal is unknown due to difficulties in defining the base of the formation in poor exposure. lithology. fineto coarse-grained cross-bedded or structureless, light coloured, mainly yellowish quartz sandstone. pebbly sandstone lags commonly rich in marine bivalves and belemnites occur locally. siltstones and heteroliths form the basal part of the formation in payer dal. boundaries. in payer dal, the lower boundary is a distinct surface separating the carbonate-cemented sandstone unit of the uppermost pelion formation from the siltstones and heteroliths of the basal payer dal formation. the upper boundary is placed where sandstones of the payer dal formation are sharply overlain by offshore siltstones and heteroliths of the bernbjerg formation. distribution. the formation occurs on kuhn ø, store koldewey, hochstetter forland and hold with hope (piasecki et al. in press; vosgerau et al. in press). geological age. the age of the formation is not wellconstrained due to a lack of ammonites. dinocyst assemblages suggest a mainly early – early late oxfordian age (table 1). subdivision. the formation is subdivided into a lower coarsening-upwards siltstone–sandstone unit and an upper sandstone-dominated unit. these two units are separated by a succession of pebbly sandstones, up to 3 m thick, and are considered informal members. facies. in payer dal, the basal part of the lower payer dal formation consists of a succession of siltstones and heteroliths which has a minimum lateral extent of a few hundred metres but cannot be traced further due to extensive scree cover (figs 16, 17). these basal sediments coarsen upwards into cross-bedded, mediumto coarse-grained sandstones which alternate with evenly laminated, fineto medium-grained sandstones or wave rippled, medium-grained sandstones (fig. 17). the sets of the cross-bedded sandstones are up to 2 m thick. the sandy foresets are commonly separated by single and double mud drapes. foreset dip azimuths are mainly towards the south-west, but bi-directionally orientated foresets also occur. a pebbly sandstone interval, up to 3 m thick, containing abundant thick-shelled bivalves and belemnites locally forms the top of the unit. 883 884 the upper unit of the payer dal formation consists mainly of fineto medium-grained sandstone and is generally less cemented than the underlying shallow marine quartz sandstones of the lower unit and the pelion formation. the best section through the upper unit is section a (figs 9a, 18; see also surlyk 1977, fig. 4) where it consists of stacked successions of trough cross-bedded, fineto medium-grained sandstones, commonly capped by oyster-rich, calcite-cemented, coarse-grained sandstones. cross-bedded sandstones with sets up to 3 m thick and foresets commonly separated by single and double mud drapes occur in the upper part of the unit in section d, south kingofjeld (fig. 18). palaeocurrent directions of the cross-beds of sections a and d are towards the s–sw. in bastians dal, the lower payer dal formation (possibly including the uppermost pelion formation) consists of a sandstone succession, c. 50 m thick, which forms a series of prominent benches separated by recessive slopes. the benches are composed of trough crossbedded, structureless and parallel laminated, calcitecemented sandstones, whereas the poorly exposed slopes appear to correspond to weakly cemented sandstones. the last sandstone bench that forms the top of the lower unit is capped by a thin persistent oyster bed overlain in many places by a thin pebble lag. the upper unit is composed of buff to yellow-red quartz sandstone; the exposed section is about 20 m thick. it shows large trough cross-sets at the base, up to 4 m thick, which grade up to small trough cross-sets in the uppermost metre (fig. 20). the sandstones are well-sorted and composed of subangular to subrounded grains at the base of the coset, with more poorly sorted and angular grains towards the top. the cross-bedded sandstone unit is fineto very coarse-grained and shows an overall coarsening-upwards trend. palaeocurrents of the cross-beds are towards the south-west. the unit is not exposed above the coset of trough cross-beds, but is probably present west of the area beneath quaternary alluvium deposits. palaeoenvironment. the payer dal formation was deposited in a tidally-influenced environment as indicated by the abundance of cross-bedded sandstones with foresets separated by single and double mud drapes and the local occurrence of herringbone cross-bedding. the cross-bedded sandstones are interpreted to represent south-westwards migrating tidal bars or sandwaves in a shallow marine embayment (surlyk 1977; surlyk & clemmensen 1983). deposition of the siltstones and heteroliths that form the base of the lower unit in the payer dal area, took place in an offshore transition to offshore environment after drowning of the shallow marine sandstones of the pelion formation. drowning was followed by shoreface progradation as reflected by the coarsening-upwards trend within the succession of siltstones, heteroliths and cross-bedded sandstones. the pebbly sandstones with bivalves and belemnites at the top of the lower unit represent a composite lag fig. 20. sandstone showing large-scale trough cross-bedding in the upper payer dal formation. bastians dal, central kuhn ø. view towards the west. deposit that is interpreted to have formed by transgressive wave ravinement of the underlying shallow marine sands. drowning resulted in a deeper water regime as reflected by the generally finer grain size of the upper unit compared to the lower unit. the coset of cross-bedded sandstones which occurs at the base of the upper unit in the bastians dal area may represent a large, composite tidal sand bar which migrated towards the south-west. bernbjerg formation facies. the bernbjerg formation covers much of the west side of kuhn ø and has been estimated to have a stratigraphic thickness of c. 450 m on kuhn ø (figs 3, 7; maync 1947; surlyk 1977; surlyk & clemmensen 1983). it is dominated by dark grey to black mudstones that are generally thin-bedded to laminated and weather to form a yellow-tan fissile mudstone. the sediments are generally rich in plant debris, ammonites, belemnites and locally the bivalve buchia (surlyk 1977). the formation overlies the payer dal formation with a sharp boundary (see definition above). it is separated from overlying rift-climax sediments of the wollaston forland group by an erosional, and locally angular, unconformity in the eastern part of the wollaston forland basin where elevated fault block crests were eroded, and by a conformity in the western down-tilted part of the block (surlyk 1977, 1978b, 1991). on kuhn ø, the age of the bernbjerg formation ranges from late oxfordian – kimmeridgian on the basis of ammonites and dinocysts. the bernbjerg formation is readily subdivided into a lower heterolithic unit, the ugpik ravine member of surlyk (2003, this volume, fig. 5), and an upper unit dominated by uniform dark mudstones. in the payer dal area, the ugpik ravine member is approximately 75 m thick (surlyk 1977, figs 4, 24). it is of (?)late oxfordian – early kimmeridgian age based on ammonites (sykes & surlyk 1976; surlyk 1977) and dinocysts. 87sr/86sr isotope values from a belemnite (table 2) found a few metres above the base of the bernbjerg formation, indicate a latest oxfordian – early kimmeridgian age (m. engkilde, personal communication 1997). in section d, south kingofjeld (fig. 4), the ugpik ravine member comprises seven coarsening-upwards heterolithic cycles, 5–14 m thick (fig. 18). the lower part of each cycle typically consists of parallel to slightly undulating heterolithic laminae and thin beds. in the upper, more coarse-grained part, wave ripple and current ripple cross-lamination may occur. the cycles are separated by sharp erosional boundaries and, in some cases, by a lag rich in belemnites, wood fragments and wellrounded quartzite pebbles up to 4 cm in diameter. the ugpik ravine member is capped by a sharp surface where heterolithic sediments are succeeded by a succession of dark grey to black mudstones. this upper unit of the bernbjerg formation is estimated to be several hundred metres thick on kuhn ø but was not measured or described in detail. palaeoenvironment. the sharp boundary at the base of the bernbjerg formation is interpreted as a major drowning surface formed during transgressive ravinement of the shallow marine sandstones of the underlying payer dal formation. the coarsening-upwards heterolithic cycles in the ugpik ravine member represent progradational events that resulted in progressive shallowing from the offshore to the lower or middle shoreface zone. the erosional boundaries and the localised lag deposits on top of the coarsening-upwards cycles are interpreted to have formed by transgressive reworking and winnowing. the sharp upper surface of the lower unit of the bernbjerg formation is interpreted as a major drowning surface marking the transition to offshore mudstones. geochemistry. hydrogen index (hi) values of six mudstone samples from the bernbjerg formation fall in the range 32–143 mg hc/g toc; toc values range between 2.8% and 5.4%. the low hi values are thought to be a consequence of the proximal setting, close to the cratonic mainland towards the west and north. most of the bernbjerg formation samples contain plant debris or coal fragments. thickness variations and sediment geometry in the payer dal region, the total thickness of the shallow marine sandstones of the pelion and payer dal formations varies from a minimum of 40 m at the crest of the fault-block to the east, to more than 350 m in payer dal itself and probably more than 500 m further west (figs 7, 18). it is unclear if the thinning towards the east is due to onlap, depositional thinning, truncation or some combination of these. however, subtle bedding plane features seen on the aerial photographs and panoramas suggest that some of the thinning is due to eastwards onlap of the formations onto the peneplaned basement or the bastians dal formation (fig. 10a). the 885 depositional onlap probably reflects differential subsidence due to the onset of fault block tilting, possibly combined with eustatic sea-level rise. in this context, it is noteworthy that some very coarse-grained, poorly sorted sandstones occur in the pelion formation or the lower unit of the payer dal formation in the east of the payer dal area, suggesting a proximal setting close to the sediment source (figs 4, 18; section e, schwarze wand). on south kuhn ø, the peneplaned surface of the basement has a regional strike of 160° and a dip of 9° to the wsw, mainly reflecting jurassic and later fault block rotation. the top of the lower unit of the payer dal formation has an average calculated regional strike of 160° and a dip of 6–7° to the wsw and represents a major drowning surface which was most likely close to horizontal at the time of deposition. bedding planes within the pelion formation and the lower unit of the payer dal formation dip 5–13° towards the wsw with an average strike of 165°. the dip variations observed in the pelion formation and the lower payer dal formation may be explained by a sigmoidal clinoform model (fig. 21). the clinoform surfaces are represented by individual dips measured on bedding planes at discrete localities. the larger of these dips (10–13°) may represent foresets of sand bars or, on a larger scale, the slopes of prograding clinoforms. the smaller values (5–7°) probably represent beds originally closer to horizontal with superimposed post-depositional structural rotation. clinoforms are difficult to define in the field, with the exception of a locality near the crest of schwarze wand, where clinoforms dip up to 15°. a sigmoidal model explains the variability of dips and fits the interpreted depositional setting with numerous progradational events separated by drowning. diagenesis the heterogeneous distribution of calcite cement and the abundance of concretions, especially in the pelion formation and the lower unit of the payer dal formation, is a striking and characteristic feature of the shallow marine sandstones. concretions range in size from a few centimetres to over 30 cm in diameter. concretionary calcite-cemented zones typically follow bedding planes, but in some places cut vertically up section. the source of the calcite cement is probably biogenic carbonate derived from the abundant calcareous shelly fauna that is associated with the cemented sandstones. the high volumes of calcite cement (34–42%) in several samples from the pelion formation and the lower payer dal formation suggest that calcite precipitation took place at shallow to moderate burial depths. concretionary zones along bedding planes are a common feature of similar age sandstones on the norwegian continental shelf, and in onshore exposures in north-west scotland and elsewhere in east greenland (fürsich 1982; walderhaug et al. 1989; bjørkum & walderhaug 1990). bjørkum & walderhaug 886 a e sw (250°) ne top lower unit payer dal fm top basement 6–7° 5–7° 5–7° 10–13° 10–13° 9° fig. 21. simplistic sigmoidal clinoform model constructed to explain structural data from south kuhn ø. the dips of the top of the lower payer dal formation and the top basement surface are regional estimates based on calculations, whereas the slope angles of the dashed surfaces are based on measured dips from bedding planes at discrete localities. the higher dip values (10–13°) probably represent foresets of sand bars or, on a larger scale, the slopes of prograding clinoforms. the lower values (5–7°) probably represent beds deposited in a sub-horizontal attitude with superimposed post-depositional structural rotation. sections a and e are c. 7 km apart; the lower payer dal formation is c. 40 m thick in section e (fig. 18). (1990) offered a detailed discussion of the nucleation and growth processes that create this fabric. porosity and permeability measurements from core plugs taken from the sandstones of the pelion and payer dal formations reflect a strongly bimodal distribution of porosity and permeability between the cemented and weakly cemented sandstones (figs 14b, c, 22). petrographic studies of sandstone samples from the upper payer dal formation (fig. 4, samples c-2, -16, -17, -39, -40, -43, -51) show that the sandstones have undergone mechanical compaction, but have not been buried deeply enough to induce silica cementation. measurements undertaken on some of the sandstones give helium porosity values of 28–32% and permeability values of 438–4900 md. the jurassic sandstones from the pelion and payer dal formations of kuhn ø are very similar to jurassic sandstones found on the mid-norwegian shelf and in the barents sea with regard to original mineralogical composition, sorting and grain size. a minor difference is that parts of the correlative fangst group of the midnorwegian shelf are coarser grained and contain less biogenic carbonate. calcite cement is not as volumetrically important in the jurassic sandstones of the midnorwegian shelf as it appears to be in samples from kuhn ø. this is probably a function of the higher content of biogenic carbonate in the kuhn ø samples. quartz cement is typically extensive in the deeply buried sandstones of the norwegian shelf, but is lacking in the samples from kuhn ø, suggesting that the jurassic deposits of kuhn ø have not been buried to depths where temperatures exceed 60–100°c (mcbride 1989). regional correlation recognition of correlative regional surfaces in the jurassic succession of the wollaston forland basin is difficult due to the abundance of local erosion surfaces in the tidal deposits and the restricted lateral extent of many of the outcrops. the most distinctive regional surfaces are the major drowning surfaces which separate the jurassic succession into a series of backstepping sedimentary units reflecting the overall middle–late jurassic transgression (surlyk & clemmensen 1983; surlyk 1991). identification and tracing of the major drowning surfaces allow cor887 100 000 10 000 1000 100 10 1.0 0.1 0.01 0 10 20 porosity (%) pe rm ea bi lit y (m d ) 30 40 fig. 22. porosity and permeability of sandstones from the pelion formation and the lower payer dal formation, kuhn ø. the bimodal distribution reflects the occurrence of cemented and poorly cemented sandstones (see fig. 14b, c). 888 sa nd sa nd sa nd sa nd sa nd sa nd ? se ct io n 1 se ct io n 2 se ct io n 3 se ct io n 4 se ct io n 5 c en tr al k uh n ø ba st ia ns d al sw w ol la st on f or la nd c ar di oc er as da l so ut he rn k uh n ø pa ye r d al so ut h n or th se ct io n 6 so ut he rn h oc hs te tt er f or la nd k ul hu s / s øn dr e m us lin ge bj er g ? ? ? ? ? ba st ia ns d al f m sh el ls c al ci te c on cr et io ns lo ca tio n of b el em ni te s co lle ct ed fo r sr is ot op ic a na ly si s (t ab le 2 ) 10 0 m be rn bj er g fm ja ko bs st ig en fm pe lio n fm pe rm ia n ca rb on at es a nd ev ap or ite s 32 k m 19 k m 31 k m m us lin ge bj er g fm d s d s d s d s d s d s c ro ss -b ed di ng tr ou gh c ro ss -b ed di ng w av e ri pp le c ro ss -la m in at io n pl an ar la m in at io n an d be dd in g st ru ct ur el es s c oa l o r ca rb on ac eo us b ed s c al ed on ia n cr ys ta lli ne b as em en t pe bb le s m aj or d ro w ni ng s ur fa ce pe lio n fm pa ye r d al f m lo w er u ni t pa ye r d al fm up pe r un it be rn bj er g fm be rn bj er g fm fi g. 2 3. c o rr el at io n o f th e ju ra ss ic s u cc es si o n o n k u h n ø w ith m o re o ff sh o re d ep o si ts t o t h e so u th i n w o lla st o n f o rl an d a n d m o re p ro xi m al , co as tn ea r d ep o si ts t o t h e n o rt h i n h o ch st et te r fo rl an d , b as ed o n b io st ra tig ra p h ic a n d 87 sr /86 sr i so to p e d at a. s ec tio n 1 is b as ed o n s u rl yk ( 19 77 ), v o sg er au ( 19 97 ) an d v o sg er au e t a l. (2 00 0) . se ct io n s 2 –5 ar e b as ed o n s u rl yk ( 19 77 ) an d t h is s tu d y; s ec tio n 2 co rr es p o n d s to s ec tio n s c an d d , s ec tio n 3 to s ec tio n a (f ig s 4, 1 8) . s ec tio n 6 is b as ed o n c le m m en se n & s u rl yk ( 19 76 ) an d s u rl yk ( 19 77 ). relation of the sedimentary units, in combination with biostratigraphic and 87sr/86sr isotope data. the major drowning surfaces are, however, difficult to recognise in both the most proximal landwards areas where nearshore sandstones occur both below and above the surfaces and in the most distal areas where the surfaces separate offshore mudstones. biostratigraphic control is limited at some levels due to the scarcity of ammonites in the sandstones and the low diversity and generally low biostratigraphic resolution of the dinocyst assemblages. kuhn ø occupies an intermediate position between the generally more offshore deposits to the south in wollaston forland and the more proximal succession to the north in hochstetter forland (fig. 6). the successions in wollaston forland and hochstetter forland are briefly described below and a correlation with the sections on kuhn ø is suggested. in the cardiocerasdal area, in south-west wollaston forland, crystalline basement is overlain by a thin permian succession of carbonates and evaporites but the contact with overlying jurassic sediments is not well-exposed. maync (1947) reported a coarse-grained lag, 3–4 m thick, rich in shells and logs occurring near the basement in cardiocerasdal. it is interpreted as a composite transgressive lag formed by shallow marine winnowing during transgression of basal fluvial deposits equivalent to the bastians dal formation. the occurrence of the ammonite kepplerites tychonis ravn at this level indicates a late bathonian age for the lag deposit (maync 1947; callomon 1993). in northern wollaston forland, the pelion formation rests directly on crystalline basement. the basal part of the pelion formation is not well exposed in the cardiocerasdal area and faulting makes thickness estimates uncertain. the uppermost 50 m of the formation are well-exposed, however, and consist of cross-bedded, tidally influenced sandstones showing s–sw palaeocurrent directions (fig. 23). calcite concretions are locally abundant, as observed in the sandstones of kuhn ø. the cross-bedded sandstones are interpreted to reflect shallow marine south-west migrating tidal bars or sandwaves (surlyk 1977; surlyk & clemmensen 1983). the pelion formation is separated from the overlying jakobsstigen formation by a major drowning surface which can be correlated with that at the top of the pelion formation on kuhn ø (fig. 23). on kuhn ø, this surface caps a regional, benchforming, carbonate-cemented sandstone and in the west kingofjeld section it is overlain by siltstones and heteroliths of the lower payer dal formation (fig. 17). the age of this major drowning surface is not well-constrained but it probably developed in late callovian time as indicated by dinoflagellate cysts from sediments below and above the surface and by a few finds of ammonites in the middle part of the jakobsstigen formation in cardiocerasdal (fig. 23). the strontium isotopic ratios from belemnites from the uppermost pelion formation in the two areas are similar (table 2) and indicate a late or possibly a middle callovian age (m. engkilde, personal communication 1997). the jakobsstigen formation is c. 130 m thick in cardiocerasdal. it is of early–middle oxfordian age, but the dating is not well-constrained. the formation consists of rhythmically interbedded coastal plain and shallow marine deposits (bojesen-koefoed et al. 1997; vosgerau et al. 2000). the jakobsstigen formation is capped by the next major drowning surface which is thought to correlate with the surface separating the lower and upper units of the payer dal formation on kuhn ø (fig. 23). the presence of coastal plain deposits within the jakobsstigen formation is indicative of deposition in a more up-dip position on the hangingwall than that represented by the cross-bedded tidal sandstones of the payer dal formation. in cardiocerasdal, the jakobsstigen formation is overlain by a strongly bioturbated, faintly wave and current rippled heterolithic unit, c. 55 m thick, deposited in the shallow offshore zone, and forming the basal part of the bernbjerg formation. ammonites indicate the upper oxfordian amoeboceras glosense chronozone (fig. 23), and 87sr/86sr isotope values from three belemnites (table 2) indicate a late oxfordian age (m. engkilde, personal communication 1997). it is capped by a major drowning surface of late amoeboceras glosense chron age; this zone extends a few metres above the drowning surface. the drowning surface is correlated with the major drowning surface separating the upper unit of the payer dal formation and the bernbjerg formation on kuhn ø (fig. 23). in cardiocerasdal, the succession above this drowning surface consists of more than 200 m of offshore mudstones of the bernbjerg formation. the jurassic succession on hochstetter forland was deposited at the head of the embayment in a more proximal setting than the succession on kuhn ø (fig. 6). it overlies upper proterozoic dolomites on the southwest side of søndre muslingebjerg, a few kilometres south-east of kulhus. coal-bearing sediments of the muslingebjerg formation occur in an isolated downfaulted block at kulhus in southern hochstetter forland (clemmensen & surlyk 1976; petersen et al. 1998). the lower boundary is not exposed but the formation is con889 sidered to be c. 20 m thick. it consists of four coal beds, up to 3.5 m thick, interbedded with lagoonal and subordinate shoreface sediments. only one coal bed occurs at søndre muslingebjerg. it is c. 0.15 m thick and is situated c. 3 m above basement. the coal beds at kulhus thus wedge out towards the south-east within a distance of a few kilometres. dinoflagellate cysts from marine sediments immediately above the uppermost coal bed at kulhus indicate the top upper callovian p. athleta chronozone (fig. 23). the coal-bearing deposits of the muslingebjerg formation are overlain by a poorly exposed succession of shallow marine sandstones, c. 100 m thick, referred to the payer dal formation. the top of the sandstone succession is dated by ammonites to the upper oxfordian, amoeboceras glosense or amoeboceras serratum chronozone (fig. 23; sykes & surlyk 1976). further north in hochstetter forland, small outcrops of lower kimmeridgian offshore mudstones have been reported (surlyk 1978a). summary and conclusions kuhn ø offers an excellent study locality of middle– upper jurassic early syn-rift stratigraphy in a half-graben setting. the jurassic succession rests directly on the peneplaned surface of the crystalline basement. the main geometrical features are a marked thickening of the jurassic from the eastern crestal area into the graben to the west and eastwards onlap onto the crystalline basement peneplain. the main conclusions are as follows. 1. five jurassic early rift formations are recognised below the rift-climax wollaston forland group on kuhn ø. they include the bastians dal (new) and muslingebjerg formations of middle jurassic age, the pelion formation (revised) of late bathonian – late callovian age, the payer dal formation (new) of early – early late oxfordian age, and the bernbjerg formation of (?)late oxfordian – kimmeridgian age. 2. the bastians dal formation consists of a succession of fluvial conglomerates and pebbly sandstones up to 150 m thick. on central kuhn ø, the coarsegrained deposits are interpreted to fill an incised valley system, probably of early and early middle jurassic age. on southern kuhn ø, the unit is thinner, possibly because the exposures are located at the margin of the valley system. the position of the incised valley system was probably controlled by zones of structural weakness. the fluvial conglomerates and sandstones were deposited during early base-level rise. during continuing and possibly accelerating base-level rise, the fluvial environments became covered with peat swamps shortly before the valley was completely filled. eventually the valley and the adjacent interfluves were flooded, probably in late bathonian – early callovian times. 3. coals of the muslingebjerg formation on south kuhn ø comprise kerogen types ii–iii, contain over 50% toc and give high hydrogen index values (up to 700 mg hc/g toc), thus indicating that they represent a good oil-prone source rock. 4. the pelion formation on kuhn ø is interpreted to reflect deposition in the lower to upper shoreface based on the dominance of evenly laminated, swaley cross-stratified, or strongly bioturbated fineto medium-grained sandstones alternating with wave rippled or trough cross-bedded, mediumto coarsegrained sandstones. 5. the payer dal formation is subdivided into two units separated by a major drowning surface. the formation is dominated by trough cross-bedded sandstones formed by migration of shallow marine tidal sandwaves and sand bars towards the s–sw. the generally finer grain size of the upper unit of the formation suggests a deeper water setting than that represented by the pelion formation and the lower unit of the payer dal formation. 6. the pelion and payer dal formations are petrographically very similar to jurassic sandstones from the mid-norwegian shelf and the barents sea with regard to original mineralogical composition, sorting and grain size. however, the pelion and payer dal formations have not been buried to sufficient depths to induce silica cementation. porosity and permeability data from the pelion formation and the lower unit of the payer dal formation show a strongly bimodal distribution between calcitecemented and weakly-cemented quartz sandstones. the calcite cement was probably derived from carbonate fossils and developed into concretionary fabrics. the upper unit of the payer dal formation is generally a weakly-cemented quartz sandstone with helium porosities around 30%. 7. the bernbjerg formation mudstones are comparable to the upper jurassic source rock of the mid890 norwegian shelf, but give low hi values (32–143 mg hc/g toc) reflecting the proximal setting relative to the cratonic mainland to the west. 8. the most distinct regional key stratigraphic surfaces are the major drowning surfaces which separate the jurassic package into a succession of backstepping sedimentary units. identification and tracing of the major drowning surfaces in combination with biostratigraphic and 87sr/86sr isotope data allow correlation of the sedimentary units within the wollaston forland basin. in proximal settings, however, shallow marine sandstones may occur both below and above the major drowning surfaces making it difficult or impossible to distinguish the sedimentary units. acknowledgements p.c. alsgaard and v.l. felt are grateful to their main sponsor amoco norway oil company who also gave them permission to publish, and to statoil and the other wollgan partners who let amoco participate in the project. we thank the following colleagues for biostratigraphic and isotope stratigraphic information: d.g. benson and s. piasecki (palynology), j.a. bergen and j.r. groves (nannoplankton and foraminifera), j.h. callomon (ammonites), and m. engkilde (sr isotopes). statoil and saga are thanked for providing dinoflagellate cyst data from the pelion, jakobsstigen, and bernbjerg formations in the cardiocerasdal area. petrography was performed by o. walderhaug and l. bonnell of rogaland research. core analysis was done by d. russel of schlumberger geoquest, and source rock analysis by g. von graas of statoil and i.l. ferriday of geolab nor. g. artigas and other amoco employees gave valuable contributions to the article. we are grateful to a.g. whitham, t. olsen and especially j.r. ineson for constructive and helpful reviews. s. broen-jensen was an 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(eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 659–722 (this volume). surlyk, f. & clemmensen, l.b. 1983: rift propagation and eustacy as controlling factors during jurassic inshore and shelf sedimentation in northern east greenland. sedimentary geology 34, 119–143. sykes, r.m. & surlyk, f. 1976: a revised ammonite zonation of the boreal oxfordian and its application in northeast greenland. lethaia 9, 421–436. vischer, a. 1943: die postdevonische tektonik von ostgrönland zwischen 74° und 75°n. br., kuhn ø, wollaston forland, clavering ø und angrenzende gebiete. meddelelser om grønland 133(1), 195 pp. vosgerau, h. 1997: depositional environments and sequence stratigraphy of a middle–upper jurassic early syn-rift succession deposited in a low gradient epeiric seaway, the wollaston forland basin, northeast greenland 1–3, 110 pp. unpublished ph.d. thesis, university of copenhagen, denmark. vosgerau, h., bojesen-koefoed, j.a., petersen, h.i. & surlyk, f. 2000: forest fires, climate and sea-level changes in a coastal plain – shallow marine succession (early–middle oxfordian jakobsstigen formation, north-east greenland). journal of sedimentary research 70, 408–418. vosgerau, h., larsen, m., piasecki, s. & therkelsen, j. in press: a new middle–upper jurassic succession of hold with hope, north-east greenland. in: stemmerik, l. & stouge, s. (eds): the jurassic of north-east greenland. geological survey of denmark and greenland bulletin. walderhaug, o., bjørkum, p.a. & nordgård bolås, h.m. 1989: correlation of calcite-cemented layers in shallow-marine sandstones of the fensfjord formation in the brage field. in: collinson, j.d. (ed.): correlation in hydrocarbon exploration, 367–375. london: graham & trotman for the norwegian petroleum society (npf). manuscript received 7 november 1996; revision accepted 27 january 1998. geological survey of denmark and greenland bulletin 31, 2014, 39-42 39 ribbed moraines formed during the retreat of the scandinavian ice sheet from eastern himmerland, ne jylland, denmark hans lerche, peter roll jakobsen and stig a. schack pedersen the glacial geology of himmerland in the north-eastern part of jylland, south of limfjorden (fig. 1) has never received any special attention. however, the occurrence of parallel ridges south of torup was mentioned by milthers (1948) who interpreted them as marginal moraines. the ridges were recently studied during mapping of eastern himmerland. systematic geological mapping of the area north and south of mariager fjord started in 2009 and was completed in 2013 (map sheet 1316 iii; pedersen et al. 2013). this was followed by the map sheet to the north (1316 iv). during the recent mapping the extent of the terrain with parallel ridges was determined (fig. 2); the western boundary is found in rold skov (pedersen & jakobsen 2005) and the eastern boundary follows an ancient coastline in eastern himmerland. the most impressive parallel ridges occur in a forested area east of madum sø where the top level of the ridge crests reaches an elevation of 95 m a.s.l. however, the majority of the crests are at 60–70 m a.s.l. and most of the ridges are c. 10–15 m high. the sediments in the ridges are dominated by coarse-grained sand and gravel, and accumulations of erratic boulders are found on the surface of the ridges. after completion of the mapping of the area, we decided to make a detailed investigation of the ridges using groundpenetrating radar (gpr) to map the internal structures of the ridges. we also studied sections in a small gravel pit at the northern boundary of siem skov (figs 3, 4). the aim of this paper is to describe the terrain with parallel ridges, which we interpret as ribbed moraines. we discuss the geological setting of the ridges in relation to the data acquired during the systematic geological mapping and the accompanying gpr survey. a glaciodynamic model for the formation of the ridges is proposed related to the recession of the ice that had its source area in south central sweden. © 2014 geus. geological survey of denmark and greenland bulletin 31, 39–42. open access: www.geus.dk/publications/bull himmerland vendsyssel 10°e limfjorden lille vildmose kongerslev rebild bakker støvring hobro aalborg torup kirke rold skov hellum skov mariager fjord fig. 2 56°50´n hadsund 5 km denmark 50 km kattegat fig. 1. map of the north-eastern part of himmerland where systematic geological mapping has been conducted over the past five years. the area with moraine ridges is indicated by a rectangle. the locations of place names mentioned in the text are shown. 0–10 10–20 20–30 30–40 40–50 50–60 60–70 70–80 80–90 90–100 2 km elevation (m a.s.l.) ridge fig. 4 madum sø fig. 2. lidar-based digital elevation model of the study area showing moraine ridges. most of the ridges are found in three separate areas. the most densely spaced and largest ridges occur in the central area. 4040 geological setting the bedrock in himmerland is dominated by maastrichtian chalk which is exposed in pits throughout the region. an erosional unconformity separates chalk from quaternary deposits. major parts of the region between mariager fjord and limfjorden are covered by glaciofluvial sand that reaches a thickness of c. 40 m at hadsund. in the small gravel pit at the northern boundary of siem skov, the succession begins with a glaciolacustrine unit, which increases in thickness towards the east. this unit is interpreted as a palaeo-kattegat deposit and may correlate with the lønstrup klint formation in vendsyssel (pedersen 2005). it is overlain by glaciofluvial sand and gravel that may correlate with the rubjerg knude formation (pedersen 2005). these proglacial deposits of the glaciodynamic sequence (according to the concept of pedersen 2012) are overlain by c. 3 m of till. the till is classified as a sandy till; it contains indicator boulders from the oslo fjord region and its fabric indicates an ice-flow direction from north to south. based on this ice movement direction the till is correlated with the kattegat till formation (houmark-nielsen & kjær 2003; pedersen 2005). the top of the till is at 56 m a.s.l. in siem skov where it forms the base of the 10–25 m high ridges. meltwater related to the swedish ice advance only played a small role in the region with the parallel ridges. to the west, glaciofluvial deposits from the swedish ice advance are found in the upper part of the rebild bakker (pedersen & jakobsen 2005). after the parallel ridges had formed, depressions were created in the landscape due to melting of bodies of stagnant ice that were left from the retreating ice front. during the early holocene valleys were eroded when the relative sea level was low. geomorphology of the parallel ridges the elongate ridges cover an area of 15 × 8 km. most of them are n–s-oriented, but there are nw–se-oriented ridges in the north-western part of the area. the ridge density is highest in the central part of the area, in hellum skov and siem skov (fig. 4) where the ridges reach elevations of 70 m a.s.l. and are c. 10 m high. the mean spacing between the ridges is 48 m in hellum skov and siem skov and the longest ridges are c. 2500 m long. towards the east and south, ridges may still be recognised but are somewhat obscured. in these areas, ridge crests are at c. 30 m a.s.l. and the ridges are 5–10 m high. georadar survey of the parallel ridges in september 2013, ground penetrating radar (gpr) surveys were carried out in hellum skov where the most pronounced ridge terrain is found. an ekko 100tm device mounted on a cart was used, and six lines perpendicular to the ridges were recorded; five lines in the northern part of the forest and one line c. 1 km to the south. transmitter voltage was 400 v and antenna frequency was centered at 100 mhz. the antennas were oriented broadside to the survey direction and separated by 1 m. traces were recorded every 20 cm and consist of stacks of eight. two representative profiles, line0007 and line0014, illustrate the internal architecture of the ridges (fig. 5). assuming a mean velocity of 0.1 m/ns, 1 m equals 20 ns. in line0014 the layers below the ridges are predominantly a west east 1 m torup kirke small gravel pit madum sø line0007 line0014 siem skovhellum skov 1 km fig. 3. exposure in a small gravel pit located c. 1.2 km ssw of torup kirke showing a section through one of the ridges. west of the anticline (a) dipping layers of glaciaofluvial sand and gravel are seen, and east of the anticline layers of flow till occur. the architecture can be compared to the structures seen in the gpr profiles (fig. 5). photograph: elina kamla. fig. 4. hill shade image of the central part of the ridge terrain. the ridges are almost parallel, mainly n–s-oriented but nw–se-oriented in the north. the locations of the two gpr profiles (fig. 5) are also shown. line0007 was acquired in a spruce plantation across relatively smooth ridges, whereas lin0014 was acquired along a gravel road crossing 11 steep-sided ridges. 41 subparallel anticlines, whereas layers below depressions are dominantly subparallel synclines. in both profiles two types of faults are found: (1) small, dominantly eastward-dipping faults and (2) larger, possibly superimposed folded thrust faults, which can be traced from near the surface down to 7–10 m below the surface (fig. 5). formation of the ridges the small e-dipping faults recognised in the gpr profiles are interpreted as thrust faults caused by pushing from the east. small w-dipping faults may indicate landslides on steep ridge flanks. large thrust faults are interpreted as listric faults with décollement surfaces c. 7–10 m below the ground. the folded lower layers of line0014 suggest ductile deformation within the ridges and the depressions. gpr data indicate that the ridges were formed by ice pushing from the east under non-permafrost conditions. we suggest that the ridges reflect a net ice recession with a mean rate of c. 50 m/year, but with a number of small advances (fig. 6). the east–west extent with ridges is almost 10 km wide, corresponding to a period of c. 200 years. a similar magnitude of recession rate of the scandinavian ice sheet has been suggested for vendsyssel, where recession from the coastal area along kattegat north of limfjorden is estimated to have occurred in the interval from 19 to 18.5 ka (sandersen et al. 2009). we suggest that shortly after termination of the last glacial maximum temperature increased rapidly, which lead to a significant recession of the ice margin. the moraines show some similarities to de geer moraines (de geer 1940; lundquist 1986; lundquist & viborg 1998; lindén & möller 2005). however, de geer moraines formed in water depths of 150–250 m, during retreat of a grounding line with a calving glacier front. in contrast, the himmerland moraines formed on dry land. west east line0007 line0014 western part line0014 western part line0014 eastern part line007 0 200 400 0 200 20 m 0 200 400 0 200 bedding fault 0 200 400 0 200 400 tw ow ay tr av el ti m e (n s) groundwater? west east fig. 5. two examples of processed ground-penetrating radar records and their interpretations (line0014 is divided into two parts). the gpr data were processed using reflex2dquick analysis. the following steps were followed prior to the interpretation of the data: (1) x-axis flip, (2) move start time to –45.2 ns, (3) normal move-out, 1 m separation, (4) dewow noise filtered, time window 10 ns, (5) bandpass frequency, 20/187 mhz, (6) topographic migration, summation width 50 traces, (7) divergence compensation gain, scaling value 1 and (8) topographic correction, square interpolation. our knowledge of radar-wave velocity variations in the survey area is limited, and hence we refrained from converting the recorded two-way travel time to depth. however, hyperbola velocity adaptions provided approximate mean velocities for each survey line. 4242 conclusions as part of the systematic mapping of eastern himmerland, a detailed investigation was made of a terrain with parallel ridges located in a 120 km2 large area centred on siem skov. the individual ridges are up to c. 2.5 km long. the ridges are 10–25 m high, and the spacing between the ridge crests varies from 50 to 100 m. the parallel ridges are interpreted as ribbed moraines that formed during recession of the ice margin at c. 19 ka. the detailed architecture of the ridges was investigated by gpr surveys which show an integrated depositional and deformational dynamic picture with narrow and steep-sided depressions filled with flow-till materials contemporaneous with fold-push and up-thrusting. references de geer, g. 1940: geochronologia suecica, principles. kungliga svenska vetenskapsakademiens handlingar ser. iii 18(6), 367 pp. houmark-nielsen, m. & kjær, k.h. 2003: southwest scandinavia, 40–15 kyr bp: palaeogeography and environmental change. journal of quaternary science 18, 769–786. lindén, m. & möller, p. 2005: marginal formation of de geer moraines and their implications to the dynamics of grounding-line recession. journal of quaternary science 20, 113–133. lundqvist, j. 1986: late weichselian glaciation and deglaciation in scandinavia. quaternary science reviews 5, 269–292. lundqvist, j. & viborg, l. 1998: isavsmältning och israndlinjer i sverige and lokaler i sverige. in: andersen, s. & pedersen, s.a.s. (eds): israndslinier i norden, 61–81 and 161–215. copenhagen: nordisk ministerråd. milthers, v. 1948: det danske istidslandskabs terrænformer og deres opstaaen. danmarks geologiske undersøgelse iii. række 28, 234 pp. (with summary in english). pedersen, s.a.s. 2005: structural analysis of the rubjerg knude glaciotectonic complex, vendsyssel, northern denmark. geological survey of denmark and greenland bulletin 8, 192 pp. pedersen, s.a.s. 2012: glaciodynamic sequence stratigraphy. in: huuse, m. et al. (eds) 2012: glaciogenic reservoirs and hydrocarbon systems. geological society (london) special publication 368, 29–51. pedersen, s.a.s. & jakobsen, p.r. 2005: geologisk kortlægning af statsskovarealerne i rold skov. systematisk geologisk kartering af statsskovarealerne i rold skov, som udgør dele af 1:25 000 kortbladene 1216 i sø og 1216 ii nø, nordlige jylland. danmarks og grønlands geologiske undersøgelse rapport 2005/81, 28 pp. pedersen, s.a.s., jakobsen, p.j., tougaard, l. & gravesen, p. 2013: geological map of denmark 1:50  000, map sheet mors, nw denmark. geological survey of denmark and greenland bulletin 28, 29–32. sandersen, p.b.e., jørgensen, f., larsen, n.k., westergaard, j.h., & auken, e. 2009: rapid tunnel-valley formation beneath the receding late weichselian ice sheet in vendsyssel, denmark. boreas 38, 834– 851. authors’ addresses h.l., department of geosciences and natural resource management, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: hans_lerche@hotmail.com p.r.j. & s.a.s.p., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. + + + + + + + + + + ++ west east ++ + + ++ a b c d basal till fig. 6. conceptual model showing the formation of moraine ridges in north-eastern himmerland. a: in winter the glacier front advances, pushing up and deforming deposits in front of the glacier. b: in summer the mass balance of the glacier is negative and the glacier front recedes. the proximal part of the deformed deposits collapses and a ridge is formed. a new outwash fan forms between the glacier margin and the ridge. c: the next winter the new outwash fan is pushed up and deformed. syntectonic deposition in the form of flow till and alluvial sediments also occurs. d: the next summer another outwash fan is formed. 51 lithological mapping using remote sensing depends, in part, on the identification of rock types by their spectral characteristics. chemical and physical properties of minerals and rocks determine their diagnostic spectral features throughout the electromagnetic spectrum. shifts in the position and changes in the shape and depth of these features can be explained by variations in chemical composition of minerals. detection of such variations is vital for discriminating minerals with similar chemical composition. compared with multispectral image data, airborne or spaceborne hyperspectral imagery offers higher spectral resolution, which makes it possible to estimate the mineral composition of the rocks under study without direct contact. arctic environments provide challenging ground for geological mapping and mineral exploration. inaccessibility commonly complicates ground surveys, and the presence of ice, vegetation and rock-encrusting lichens hinders remote sensing surveys. this study  addresses the following objectives: 1. modelling the impact of lichen on the spectra of the rock substrate; 2. identification of a robust lichen index for the deconvolution of lichen and rock mixtures and 3. multiscale hyperspectral analysis of lithologies in areas with abundant lichens. modelling the impact of lichen cover spectral mixing of lichens and bare rock can shift the wavelength positions of characteristic absorption features and complicate the spectral mapping of minerals and lithologies. salehi et al. (2017) investigated how surficial lichen cover affects the characteristics of shortwave infrared mineral absorption features and the efficiency of automated extraction of absorption features. for this purpose, mixed spectra were synthetically generated from laboratory spectra of common rock-forming minerals and lichens. wavelength displacements of characteristic absorption features for each mixed spectrum were then analysed as a function of lichen cover percentage (see an example in fig. 1). by quantifying lichen hyperspectral analysis of lithologies in the arctic in areas with abundant lichen cover sara salehi 2000 2100 2200 2300 2400 wavelength (nm) 0.15 0.02 0.25 0.30 0.35 0.40 re fle ct an ce kimberlite lichen hull a b quotient 2240 2260 2280 2300 2320 2340 wavelength (nm) 0.85 0.90 0.95 1.00 h ul l q uo tie nt 0 kimberlite / 100 lichen 50 kimberlite / 50 lichen 100 kimberlite / 0 lichen fig. 1. a: averaged spectra of pure rock and lichen for a kimberlite substrate in shortwave infrared range. b: the corresponding hull quotient (clark & roush 1984) and band centres of mixed spectra associated with the antigorite absorption feature. the 10% spectral intervals used to investigate the main absorption features are highlighted. x: wavelength positions of local minima (salehi et al. 2017). © 2018 geus. geological survey of denmark and greenland bulletin 41, 51–55. open access: www.geus.dk/bulletin http://www.geus.dk/bulletin 5252 cover effects on mineral absorption features, this study highlights the importance cautious interpretation in areas characterised by abundant, lichen-covered outcrops. this can be of significant importance for mineral and deposit identification, because slightly shifted features for a given spectrum caused by lichen cover can be erroneously identified as a path to a deposit. salehi et al. (2017) showed that spectral shifts caused by lichens are not constant, i.e. each mineral spectral feature may be affected differently depending on the shape of the lichen spectrum. for example, the absorption feature related to the chlorite mineral group around 2254 nm is shifted towards longer wavelength, while the one around 2320 nm is shifted towards shorter wavelength and the 2380 nm band maintains its spectral characteristics. spectral shifts are not only related to rock/lichen proportions but also to the modal abundance of minerals in certain rock types. background minerals and associated overlapping features will have an effect on the related absorption depth and play a critical role in the scale of wavelength displacement. identification of a robust lichen index the ability to distinguish a lichen cover from its rock/mineral substrate is important, and decomposition of a mixed pixel into a collection of pure reflectance spectra can improve the use of hyperspectral methods for mineral exploration. in order to identify spectral indices that can directly reflect the ratio of the rock and lichen in hyperspectral data, a number of index structures were assigned to an optimisation algorithm, which was tasked to find the best values for the location of the bands along the reflectance spectra measured in the laboratory (salehi et al. 2016). in order to further investigate the functionality of the indices for the airborne platform, the spectra were resampled to hymap resolution. the indices proposed by salehi et al. (2016) proved robust to the type of the substrate rock and permitted an estimate of the lichen cover with acceptable, albeit varying, levels of error. the results revealed that the ratio between r894-1246 and r1110 explains most of the variability in the hyperspectral data at the original laboratory resolution (r2=0.769). however, the normalised index incorporating r1106-1121 and r904-1251 yields the best results for the hymap resolution (r2=0.765). re fle ct an ce (o ffs et fo r c lar ity ) wavelength (nm) wavelength (nm) olivine (usgs) talc (usgs) anthophyllite (usgs) hornblende (usgs) hornblende (hymap) antigorite-anthophyllite (hymap) talc-olivine/pyroxene (hymap) actinolite (hymap) serpentine-olivine (hymap) serpentine (usgs) actinolite (usgs) actinolite-hornblende (hymap) a b 500 1000 1500 2000 2200 2300 2400 fig. 2. spectra of extracted end members compared with selected minerals from the usgs spectral library in envi software. a: full spectral range. b: shortwave infrared range. 53 the proposed methodology has the advantage of not requiring a priori knowledge about the exact effects of lichens – or any other substance – on the reflectance of the mixtures. instead, this information is obtained by an automated trial and error process. therefore, this technique can also be beneficial for identification of sensitive bands and indices for deconvolution of any other mixed spectra, whether synthetic as in this case, or obtained directly from the samples. multiscale hyperspectral analysis of lithologies with abundant lichen cover two sets of hyperspectral data acquired by airborne hymap (350–2500 nm) and light-weight rikola (500–900 nm) sensors were chosen to investigate the potential of visible near infrared and shortwave infrared spectral range for detailed lithological mapping in the nagssugtoqidian orogen of west greenland, where an ultramafic rock unit with abundant lichen cover is exposed. the extent to which geological information derived from airborne data is retained in the rikola serpentine + pyroxene 0 end-member abundance (%) serpentine + olivine/pyroxene amphibole (actinolite/hornblende) lichen vegetation 0.25 0.50 0.75 1 0 0.25 0.50 0.75 1 0 0.25 0.50 0.75 1 0 0.25 0.50 0.75 1 0 0.25 0.50 0.75 1 66 °4 4' 38 ''n 66 °4 4' 10 ''n 52°31'26'' w 52°30'22'' w 500 m fig. 3. the result of unmixing analysis and the abundance of mafic-ultramafic minerals using hymap data. masked pixels are indicated by black colour. ricola vnir b a b c d fig. 4. a: supracrustal rocks in the innarsuaq region comprising a kilometre-sized body of mafic-ultramafic, looking north. b: bright green and black amphibole and biotite at the corner (alteration zone). c: host ultramafic rock; green amphibole to the left and white talc vein in the middle. d: 50 cm long asbestos fibres. 5454 hyperspectral data, is examined as an insight to future dronebased hyperspectral mapping capabilities and the possibility of extracting valuable mineralogical and lithological information using such platforms. the airborne hyperspectral dataset is corrected for abnormal pixels and removal of bad bands (such as water vapour absorption features and noisy bands) prior to atmospheric correction. dark pixels, snow, clouds and water were filtered out. next, the spatial–spectral end-member extraction method (rogge et al. 2007) is used to derive an image end-member set. this makes an assessment of subtle lithological variability across a given study area possible. these end members are then sorted based on expert knowledge of known spectral features (water, snow, vegetation, lichen and geological materials) followed by a more detailed sorting into individual classes within each category. subtle shortwave infrared features related to key minerals in the geological materials are particularly important. the resulting sorted end-member classes are subsequently averaged to produce a final end-member set. a final set of six geological end members (fig. 2), and two end members related to vegetation and lichens are deducted from expert-based analysis. figure 2 mafic lichen and vegetation 500 m 50 m ultramafic a b c w e 0 0.25 0.50 0.75 1 0 0.25 0.50 0.75 1 0 0.25 0.50 0.75 1 0 0.25 0.50 0.75 1 0 0.25 0.50 0.75 1 actinolite serpentine/olivine talc anthophyllite lichen and vegetation legend (c). end-member abundance (%) fig. 5. a: true-colour hyperspectral image mosaic generated using the rikola camera. b: minimum noise fraction false-colour image: red: band 6. green: band 2. blue: band 1. c: result of spectral unmixing analysis and the abundance of mafic-ultramafic minerals. 55 shows a plot of the extracted end members using the spatial– spectral end-member extraction method and the corresponding signatures from the united states geological survey (usgs) spectral library (kokaly et al. 2017). the shortwave infrared spectral characteristics of the ultramafic rocks studied here were controlled by amphibole minerals as exemplified by hornblende, actinolite and anthophyllite (fig. 3). the absorption features in the shortwave infrared region are located at 2320 and 2380 nm and are of the same order of magnitude. the shortwave infrared spectrum of olivine-rich rocks clearly reflects a mixture of antigorite serpentine with a characteristic stronger absorption feature at 2320 nm. a less distinct absorption feature at 2310 nm is present for rocks enriched in talc. fractional abundances of the end members within the scene are determined using an iterative implementation of spectral mixture analysis method (rogge et al. 2007). the interpretation of hymap data revealed a number of mafic and ultramafic complexes in the border area between the parautochthonous and allochthonous zones of the nagssugtoqidian orogen. one such complex occurs to the east of the head of the fjord kangerluarsuk, here referred to as innarsuaq (see fig. 1 of salehi & thaarup 2018, this volume). as can be seen from fig. 3, the predictive map from the innarsuaq area displays a complex distribution of exposed bedrock, a feature confirmed during a brief field visit. the results were validated using expert knowledge of spectral characteristics of lichens and mineralogy, as well as spectral measurements of field samples and associated xrd results. the rikola camera was operated in ground-based mode and panned stepwise to acquire a set of five overlapping images. the images were corrected for geometric, radiometric and topographic effects and stitched to a continuous mosaic (figs 4, 5). the distribution of lithological units were then mapped using the minimum noise fraction method (kruse et al. 1993).the information regarding mineral abundances were retrieved using the spectra unmixing procedure (fig. 5). conclusions 1. lichen effects on the spectra of their rock substrate have important implications for the geological analysis of airborne/spaceborne hyperspectral data where rock-encrusting lichens partially obscure exposed bedrock. 2. analysis of airborne hyperspectral data can result in highquality regional mapping products capable of discriminating geological materials of interest based on subtle spectral differences. the map product generated from the rikola scenes in this study captures the broad geological patterns and many of the lithologies generated from the airborne data, although some spectral and lithological discrimination is lost due to its more limited wavelength range. 3. the performance of hyperspectral data acquired from different platforms and at various scales is investigated for qualitative mapping of arctic mineral resources in the presence of abundant lichens. the application of such technologies to extract detailed geological information from complex inaccessible regions of greenland certainly has a very low cost/benefit ratio in comparison to traditional geological fieldwork. future space-borne hyperspectral sensors will offer new possibilities to expand the scale of mapping in greenland. integration with other remote sensing datasets such as magnetic data will simplify mineral exploration and geological mapping in the arctic. acknowledgments the helmholtz institute freiberg is thanked for the use of rikola hyperspectral imager. references clark, r.n. & roush, t.l. 1984: reflectance spectroscopy: quantitative analysis techniques for remote sensing applications. journal of geophysical research: solid earth 89(b7), 6329–6340. kokaly, r.f., clark, r.n., swayze, g.a., livo, k.e., hoefen, t.m., pearson, n.c., wise, r.a., benzel, w.m., lowers, h.a., driscoll, r.l.& klein, a.j. 2017: usgs spectral library version 7. u.s. geological survey data series 1035, 61 pp., http://dx.doi.org/10.3133/ds1035 kruse, f.a., lefkoff, a., boardman, j., heidebrecht, k., shapiro, a., barloon, p. & goetz, a. 1993: the spectral image processing system (sips) – interactive visualization and analysis of imaging spectrometer data. remote sensing of environment 44, 145–163 rogge, d.m., rivard, b., zhang, j., sanchez, a., harris, j., & feng, j. 2007: integration of spatial–spectral information for the improved extraction of endmembers. remote sensing of environment 110, 287–303. salehi, s. & thaarup, s. 2018: mineral mapping by hyperspectral remote sensing in west greenland using airborne, ship-based and terrestrial platforms. geological survey of denmark and greenland bulletin 41, 47–50 (this volume). salehi, s., karami, m. & fensholt, r. 2016: identification of a robust lichen index for the deconvolution of lichen and rock mixtures using pattern search algorithm (case study: greenland). international archives of the photogrammetry, remote sensing & spatial information sciences xli-b7, 973–979. salehi, s., rogge, d., rivard, b., heincke, b.h. & fensholt, r. 2017: modeling and assessment of wavelength displacements of characteristic absorption features of common rock forming minerals encrusted by lichens. remote sensing of environment 199, 78–92. author’s address s.s., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ssal@geus.dk. http://dx.doi.org/10.3133/ds1035 mailto:ssal@geus.dk geological survey of denmark and greenland bulletin 38, 2017, 49-52 49 the onshore nuussuaq basin in west greenland is important for hydrocarbon exploration since many of the key petroleum systems components are well exposed and accessible for study. the basin has thus long served as an analogue for offshore exploration. the discovery of oil seeps on disko, nuussuaq, ubekendt ejland, and svartenhuk halvø (fig. 1) in the early 1990s resulted in exploration onshore as well. in several wells, oil stains were observed in both the siliciclastic sandstone and in the volcanic series. an important aspect of any petroleum system is a high quality reservoir rock. the aim of this paper is to review petrophysical aspects of the reservoir potential of key stratigraphic intervals within the nuussuaq and west greenland basalt groups. reservoir parameters and porosity–permeability trends for potential siliciclastic and volcanic reservoirs within the relevant formations of the nuussuaq basin are discussed below. geological setting the nuussuaq basin formed in the cretaceous–palaeogene as part of a complex system of linked rift basins that developed along west greenland during the opening of the labrador sea and baffin bay (oakey & chalmers 2012). as a result of neogene uplift, the sediments and overlying palaeogene volcanic rocks are exposed on disko, nuussuaq, upernivik ø and svartenhuk halvø (fig. 1, dam et al. 2009). the sedimentary succession is interpreted as deposited in fluvial, delta, shelf and deep marine environments, and is divided into ten formations forming the nuussuaq group (fig. 2). the overlying west greenland basalt group (wgbg) includes subaerial lava flows and hyaloclastite breccias (larsen et al. 2016). a companion paper to this contribution (sørensen et al. 2017, this volume) provides additional information regarding the structural development and potential source rock distribution of the nuussuaq basin. potential hydrocarbon reservoirs of albian–paleocene age in the nuussuaq basin, west greenland morten l. hjuler, niels h. schovsbo, gunver k. pedersen and john r. hopper halvø ubekendt ejland uummannaq d i s k o n u u s s u a q upernivik ø ik kq it gro gw gt fp93 ma um fp94 ak gk ge pk an ki qi ikka rk uk ak pi well outcrop nuussuaq group quarternary units basement ice sheet fault oil shows west greenland basalt group formationoutcrop member ki ka atane kangilia kingittoq kangilia kingittoq annertuneq congl. pi ak ik ggu247801 ikorfat fault 566 m atane cores formation logswell faults atanepingu skansen ak fp93 it fp93-3-1 itilli fault zone 139 m atane kangiliaataata kuua qilakitsoq fp94 kq fp94-11-04 kuugannguaq-qunnilik fault 340 m itilli qi rk ge gane-1/1a 631 m agatdal+vaigat core logs atane atane qilakitsoq ravn kløft qilakitsoq ravn kløft+kingittoq gk gank-1 364 m kangilia+vaigat gro gro-3 itilli+kangilia+ well logs pk gt gant-1 891 m itilli+kangilia itillipingunnguup kuua anariartorfik an ik gw ganw-1 199 m vaigat agatdal+vaigat itilli atane anariartorfik ikorfat anariartorfik ravn kløft+kingittoq uk ma marraat-1 448 m vaigat well logs itilliukalersalik anariartorfik um umiivik-1 1200 m itilli+kangilia greenland 20 km20 km 54°w 71 ° 70 °3 0' 70 °n 71 °3 0' n svartenhuk 52°w eastern limit of n uussuaq basin eastern limit of n uussuaq basin fig. 1. geological map of the study area showing well and outcrop locations. abbreviations of outcrop, well and fault names are explained to the right, where lithostratigraphic units occurring at outcrops and wells are also listed. © 2017 geus. geological survey of denmark and greenland bulletin 38, 49–52. open access: www.geus.dk/publications/bull 5050 lithology and distribution of relevant formations the atane formation is known from the eastern part of the nuussuaq basin east of the kuugannguaq–qunnilik fault, (kq fault, figs 1, 2). the formation is up to 800 m thick in individual outcrops and consists of delta deposits, which include laterally extensive sandstone sheets (dam et al. 2009). the mudstone-dominated, marine itilli formation (fig. 2) is known from northern and western nuussuaq (west of the ikorfat fault) and is more than 2.5 km thick (sønderholm & dam 1998). the umiivik member crops out in northern nuussuaq between the ikorfat and the itilli faults (fig. 1), whereas the anariartorfik member crops out west of the kq fault. the up to 438 m thick marine kangilia formation (fig. 2) crops out on northern nuussuaq between the ikorfat and itilli faults (fig. 1) and has been drilled in the gant-1 and gro-3 wells. it has also been measured in an outcrop on southern nuussuaq. mudstones dominate in the outcrops while sandstones dominate in the wells. the annertuneq conglomerate member mainly comprises conglomerates and sandstones (dam et al. 2009). the sub-marine to marine agatdal formation (fig. 2) is known from the agatdal area on central nuussuaq west of the ikorfat fault as well as from the gro-3 well west of the k–q fault (fig. 1). the formation is up to 148 m thick and consists of mudstones, sandstones and conglomerates (dam et al. 2009). the volcanic vaigat formation (fig. 2) is up to 1600 m thick in western nuussuaq and northern disko and increases to at least 5 km in thickness on ubekendt ejland (larsen et al. 2016). the lower part of the formation is dominated by hyaloclastic breccias that are overlain by lava flows. methods ten wells and eight sedimentary outcrop successions located in the nuussuaq basin and described by sønderholm & dam (1998) and dam et al. (2009) were analysed to quantify the sandstone component of the siliciclastic formations of the nuussuaq group. all sandstone intervals were defined as potential sandstone reservoirs and expressed as the cumulative sandstone thickness of a formation. potential reservoir content was defined as the cumulative sandstone thickness divided by formation thickness. the cumulative sandstone thickness in outcrops and wells was estimated from sedimentological logs and core descriptions. mudstones and heteroliths were classified as non-reservoir lithologies and included in the term ‘shale’. in the uncored gro-3 well, sandstone content was determined from wire-line logs and implies a shale content of 0–15%. five wells containing volcanic successions of the wgbg were analysed to identify potential reservoir sections within the hyaloclastite successions. these were quantified as the cumulative thickness of hyaloclastite breccias. lava flows were classified as non-reservoirs. porosity–permeability trends were established for the gane-1/1a, gant-1 and marraat-1 wells based on core analysis data. the porosity of the clastic formations in gro-3 was calculated from the density wire-line log. sandstones with porosities >10% are referred to as porous sandstones. a porosity log of the gane-1/1a well was generated from gamma-ray and density logs obtained from core scans. maligât u pe rn ivi k n æs slibestensfjeldet kome atane itilli agatdal vaigat kangilia eqalulik atanikerluk wgbg quikavsak n uu ss ua q g ro up al bi an dan. pa le oc en e u pp er c re ta ce ou s lo w er c re ta ce ou s cen. tur. maas. se lan di an cam. san. con. w el l o ut cr opseries stage group formation member skansen ravn kløft kingittoq qilakitsoq aaffarsuaq anaanaa naujánguit ordlingassoq annertuneq conglomerate anariartorfik umiivik kussinerujuk fig. 2. lithostratigraphic scheme of the nuussuaq group and the lowermost part of the west greenland basalt group (wgbg). members studied here are indicated to the far right. modified from dam et al. (2009). 0 100 200 300 400 500 c um ul at ive s an ds to ne th ick ne ss (m ) 0 500 1000 1500 2000 formation thickness (m) 10% sandstone 50 % sa nd sto ne 10 0% sa nd sto ne 25% san dstone agatdal fm kangilia fm itilli fm atane fm fig. 3. formation thickness versus cumulative sandstone thickness, data from the analysed wells and outcrops in the nuussuaq basin. 51 reservoir properties of formations containing potential hydrocarbon reservoirs potential reservoirs were identified within sandstones of the atane, itilli, kangilia, and agatdal formations and hyaloclastite breccias of the vaigat formation (fig. 2). the main reservoir parameters including cumulative reservoir thickness (crt), potential reservoir content (prc) and porosity–permeability data are summarised below, with details in the electronic supplement (es). atane formation. crt is in the range 26–360 m corresponding to a prc of 45–80% (fig. 3 and es). the reservoir quality is well developed in central nuussuaq with sandstone porosities of 5‒25% (mean 17%) and air permeabilities of 0.5‒150 md (appel & joensen 2014). itilli formation. crt is in the range 21–458 m corresponding to a prc of 2–51% (fig. 3 and es). the relatively low content of potential sandstone reservoirs compared to the atane, kangilia and agatdal formations reflects significant variations within and between the members of the itilli formation. in the gro-3 well, the individual sandstone units are up to 100 m thick and average porosity is 4%. kangilia formation. crt is in the range 0–134 m corresponding to a prc of 0–72% indicating significant lateral variation in depositional environments (fig. 3 and es). in the mudstone-dominated successions, the annertuneq conglomerate member constitutes a potential reservoir. in the gro-3 well, porosities range between 5–17% (average 6%). agatdal formation. crt is in the range 71–148 m corresponding to a prc of 50–67% (fig. 3 and es). porosities of 6–21% and permeabilities up to 9 md were measured in the gane-1 well (fig. 4), and in the gro-3 well, an average formation porosity of 9% was assessed. vaigat formation. several oil shows have been encountered in the hyaloclastites and lava flows. only the hyaloclastites seem to possess the necessary permeability to constitute a reservoir despite indications of lower average porosities (6%) than the lava flows (8%). for identical porosity values, the permeability is 30 times higher for the hyaloclastites compared to the lava flows (fig. 4). the crt of the hyaloclastites is in the range 80–671 m corresponding to a prc of 84–100% (fig. 2 and es). porosity–permeability relations of potential reservoirs the regional distribution of porosity and permeability in the nuussuaq basin deposits is poorly known due to the scarcity of core measurements or studies concerned with the effects of diagenesis on reservoir quality (kierkegaard 1998). one relatively swift way of improving the porosity database is by using core scans to generate a porosity log (pedersen et al. 2013). this is shown for the agatdal formation in gane-1/1a in fig. 5. the core log-generated 0.01 0.1 1 10 100 ai r p er m ea bi lit y (m d ) 0 5 10 15 20 he porosity (%) trend a trend b marraat-1 wells formations gane-1/1a gant-1 agatdal fm vaigat fm (lava flows) kangilia fm vaigat fm (hyaloclastite) itilli fm oil shows gas fig. 4. helium porosity and air permeability for selected wells. oil shows are abundant along trend a (hyaloclastite reservoirs). both oil and gas shows occur in the lava flows (trend b). the siliciclastic samples generally plot between the two trends. oil gas 600 650 700 depth (m) porosity (%) air permeability (md) core gamma ray (api) 0 5 10 10 0 200 400 600 80010.10.012015 sandstone shale intrusive rock core analysis core log-porosity core gamma ray gane-1/1a gane-1 gane-1a lithologyshowscore and log measurements, agatdal fm fig. 5. stratigraphical variation of lithology, porosity and permeability within the agatdal formation in the gane-1/1a well. 5252 porosity curve fits nicely with the core measurements and the core log-derived lithology corresponds well with the lithological log from dam et al. (2009). core analysis data from the marraat-1, gane-1/1a, and gant-1 wells were used to outline porosity–permeability trends within the siliciclastic and volcanic successions (fig. 4). in the siliciclastic reservoirs, the porosity and permeability data show a high degree of scatter and no consistent trend can be identified. the large range in porosity and permeability values probably reflects variations in grain size and diagenesis. a large part of the porosity in the gant-1 sandstones is secondary and related to dissolution of detrital feldspar grains (kierkegaard 1998). the hyaloclastite samples show a relatively high permeability–porosity ratio (trend a in fig. 4) compared to the lava flows (trend b in fig. 4). at 10% porosity, the expected permeability is 5 md for hyaloclastites, but only 0.8 md for lava flows, a tendency assumed to reflect textural control on permeability. in hyaloclastites, a network of connected pores ensures fluid or gas flow, whereas the isolated pore systems in lava flows strongly impede permeability, even at high porosity. conclusions the onshore nuussuaq basin in west greenland contains potential hydrocarbon reservoirs within the siliciclastic atane, itilli, kangilia and agatdal formations, and within the hyaloclastite intervals of the vaigat formation. the siliciclastic reservoirs occur in a wide range of geological environments from fluvial over deltaic to slope and marine settings. the potential reservoir sandstone content is generally more than 50% for the sections studied from the atane, kangilia and agatdal formations, but significantly lower in the itilli formation. the cumulative sandstone thickness is mostly >100 m for all formations, including the itilli formation. porosity and permeability data suggest that sandstone and hyaloclastite reservoirs may be of good quality with porosities up to 20%. permeabilities are mostly below 10 md. however, porosity and permeability data are restricted to the western part of nuussuaq and the diagenetic control on the reservoir quality is poorly understood regionally. acknowledgements this contribution is partly the result of a project funded by the ministry of mineral resources, greenland. references appel, a.u. & joensen, i.á. 2014: prograderende deltaaflejringer fra øvre kridt, atane formationen, nuussuaqbassinet, centrale vestgrønland. unpublished bachelor thesis, department of geosciences and natural resource management, university of copenhagen, 49 pp. dam, g., pedersen, g.k., sønderholm, m., midtgaard, h., larsen, l.m., nøhr-hansen, h. & pedersen, a.k. 2009: lithostratigraphy of the cretaceous–paleocene nuussuaq group, nuussuaq basin, west greenland. geological survey of denmark and greenland bulletin 19, 171 pp. kierkegaard, t. 1998: diagenesis and reservoir properties of campanian – paleocene sandstones in the gant#1 well, western nuussuaq, central west greenland. geology of greenland survey bulletin 180, 31–34. larsen, l.m., pedersen, a.k., tegner, c., duncan, r.a., hald, n. & larsen, j.g. 2016: age of tertiary volcanic rocks on the west greenland continental margin: volcanic evolution and event correlation to other parts of the north atlantic igneous province. geological magazine 153(3), 487–511. oakey, g.n. & chalmers, j.a. 2012: a new model for the paleogene motion of greenland relative to north america: plate reconstruction of the davis strait and nares strait regions between canada and greenland. journal of geophysical research 117, 1–28. pedersen, g.k., schovsbo, n.h. & nøhr-hansen, h. 2013: calibration of spectral gamma-ray logs to deltaic sedimentary facies from the cretaceous atane formation, nuussuaq basin, west greenland. geological survey of denmark and greenland bulletin 28, 65–68. sønderholm, m. & dam, g. 1998: reservoir characterisation of western nuussuaq, central west greenland. danmarks og grønlands geologiske undersøgelse rapport 1998/6, 36 pp. sørensen, e.v., hopper, j.r., pedersen, g.k., nøhr-hansen, h., guarnieri, p., pedersen, a.k. & christiansen, f.g. 2016: inversion structures as potential petroleum exploration targets on nuussuaq and northern disko, onshore west greenland. geological survey of denmark and greenland bulletin 38, 45–48. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen, k, denmark. e-mail: mlh@geus.dk geological survey of denmark and greenland bulletin 33, 2015, 85-88 85© 2015 geus. geological survey of denmark and greenland bulletin 33, 85–88. open access: www.geus.dk/publications/bull reserves and resources for co2 storage in europe: the co2stop project niels poulsen, andrei bocin-dumitriu, sam holloway, karen kirk, filip neele and nichola smith th e challenge of climate change demands reduction in global co2 emissions. in order to fi ght global warming many countries are looking at technological solutions to keep the release of co2 into the atmosphere under control. one of the most promising techniques is carbon dioxide capture and storage (ccs), also known as co2 geological storage. ccs can reduce the world’s total co2 release by about one quarter by 2050 (iea 2008, 2013; metz et al. 2005). ccs usually involves a series of steps: (1) separation of the co2 from the gases produced by large power plants or other point sources, (2) compression of the co2 into supercritical fl uid, (3) transportation to a storage location and (4) injecting it into deep underground geological formations. co2stop is an acronym for the co2 storage potential in europe project. th e co2stop project which started in january 2012 and ended in october 2014 included data from 27 countries (fig. 1). th e data necessary to assess potential locations of co2 storage resources are found in a database set up in the project. a data analysis system was developed to analyse the complex data in the database, as well as a geographical information system (gis) that can display the location of potential geological storage formations, individual units of assessment within the formations and any further subdivisions (daughter units, such as hydrocarbon reservoirs or potential structural traps in saline aquifers). finally, formulae have been developed to calculate the storage resources. th e database is housed at the joint research centre, the european commission in petten, the netherlands. background and methods co2 storage resource assessment a resource can be defi ned as anything potentially available and useful to man. th e pore space in deeply buried reservoir rocks that can trap co2 is a resource that can be used for co2 storage. it is of utmost importance to be aware that the mere presence of a resource does not indicate that any part of it can be economically exploited, now or in the future. a reserve can be defi ned as that part of a resource that is available to be economically exploited now using currently available technology. th us, in order to move from a resource estimate to a reserve estimate, a whole series of technical, economic, legal and socio-economic criteria must be applied. th ese criteria will then identify the fraction of the resource that can actually be economically exploited in a particular jurisdiction area, using available technology. consequently, a very high level of technical assessment is required to demonstrate the existence of a co2 storage reserve, and in most cases these kinds of resources are only available within a demonstration or commercial storage project. for these reasons, it was impossible to defi ne any co2 storage reserves in the present project. 16°w 16°e 32°e 48°e 56°e 64°e 50°n 42°n 0 16°e0° 58°n 50°n 500 km 42°n 66°n co 2 stop project covered by geological surveys universities national institutes geol. survey/university not in co 2 stop latvia covered by the estonian-latvian transboundary project 8°e fig. 1. twenty-seven countries participated in the co2stop project. latvia was covered by the estonian–latvian border project. the following member states of the european union participated: austria, belgium, bulgaria, croatia, czech republic, denmark, estonia, france, germany, greece, hungary, ireland, italy, latvia, lithuania, the netherlands, poland, portugal, romania, slovakia, slovenia, spain and uk and the following non-member states: macedonia, norway, serbia and switzerland. 8686 storage mechanisms co2 can be retained in reservoir rocks by a number of mechanisms: (1) structural and stratigraphic trapping, in which co2 is retained by impermeable barriers, (2) residual trapping, in which free phase co2 is trapped by capillary forces in pore spaces, (3) dissolution of co2 into pore fl uids, (4) precipitation of co2 into minerals and (5) adsorption onto shale or coal layers. only the fi rst two of these mechanisms are signifi cant within a co2 storage project’s time frame of 10 to 50 years; the other mechanisms take much longer (van der meer & van wees 2006). th erefore, most previous studies of co2 storage resources (e.g. usgs assessment; doe storage atlases; norwegian assessment; iea best practices document) focussed on determining the amount of co2 that can be retained in conventional reservoir rocks as a dense fl uid in the fl uid-fi lled pore spaces between the grains that make up the matrix of the rock and in fl uid-fi lled fractures. moreover, the vast majority of the co2 will be trapped either in structural and stratigraphic traps or by capillary forces as a residual saturation (bachu et al. 2007). constraints on co2 storage capacity each jurisdiction area contains a given amount of pore space within its subsurface. th e total resource of pore space that is potentially available for co2 storage is that part which can be fi lled with, and will retain, injected co2. geology and physics dictate that this will be far less than the available total pore space. th ese limitations mean that only a small fraction of the total resource of pore space can be fi lled with co2. it is possible to defi ne a common method that can be used to estimate the fraction of the total pore space resource that can be used for storage (brennan 2014). if appropriate co2 densities at reservoir conditions are applied to this volume, this allows estimation of the theoretical co2 storage resource. in practice, only a fraction of the theoretical co2 storage resource in any given jurisdiction area can actually be utilised – for a variety of technical, economic, legal and social reasons. in the co2stop project, the pore space in a jurisdiction area is subdivided into reservoir formations. th ese are mappable bodies of rock which display mainly suffi cient porosity and permeability. each reservoir formation contains one or more storage units. a storage unit is defi ned as a part of a reservoir formation that is found at depths greater than 800 m and which is covered by an eff ective cap rock. th ese units are potential co2 storage units and they form the basis for the co2 storage assessments made in the co2stop project. each storage unit may contain one or more daughter units. daughter units are defi ned as structural or stratigraphic traps which have the potential to immobilise co2 within them, e.g. structural domes or proven oil and gas fi elds. th e storage potential of daughter units can be estimated separately in co2stop. the co2stop method th e co2stop project has established a database, a geographical information system (gis; esri’s arcgis 10) and a calculation engine that can provide probabilistic estimates of co2 storage capacities. th e data analysis & interrogation tool is a combination of microsoft access (data interrogation tool), and excel (storefit tool) with external code (linked to excel) to perform injection rate calculations. calculations carried out with the database analysis & interrogation tool include: storage capacity, injection rates and stochastic analyses of the storage capacity and injection rates (fig. 2). th e work to establish internationally recognised standards for capacity assessments was initiated by the carbon sequestration leadership forum (cslf) about a year before the start of the european union geocapacity project, and a cslf task force has been active since. th e paper ‘estimation of co2 storage capacity in geological media – phase 2’ by bachu et al. (2007) published by the cslf presents comprehensive defi nitions, concepts and methods to be used in estimating co2 storage capacity. as in the eu geocapacity, the co2stop method complies with the cslf recommendations. th e methods and calculations for determining the fractions of the resource, used in the co2stop project, also align with the recent international energy agency proposals for harmonising co2 storage capacity estimation methods (heidug 2013). th e co2stop method estimates the tasr (see below) and the storage resource in structural and stratigraphic traps, which have later been divided into two subsets: hydrocarbon fi elds and aquifer daughter units. the technically accessible co2 storage resource (tasr) th e co2stop calculation engine can produce a resource estimate that is similar to the technically accessible co2 storage resource (tasr) estimated by the us geological survey (brennan et al. 2010; blondes et al. 2013; u.s. geological survey geologic carbon dioxide storage resources assessment team 2013). th is is the fraction of the theoretical storage resource that can be accessed using all currently available technologies regardless of cost. th e international energy agency recommended that the fi rst step in all co2 storage resource estimates should be to assess the tasr (heidug 2013). 87 th e co2stop estimate diff ers in one main respect from the tasr estimated by the u.s. geological survey method, namely that co2stop adds the storage capacity of hydrocarbon fi elds to that of the saline aquifers. th is has to be done because the pore volume of the hydrocarbon fi elds is not provided in the project’s database, so it cannot be subtracted from the pore volume of the storage units before their storage capacity is estimated. th ere are other minor diff erences in the constraints and assumptions; nevertheless, the two methods produce results that are suffi ciently similar to allow them to be compared. results th e assessment of the various fractions of the co2 geological storage resource performed in the co2stop project is currently only at a provisional level. unfortunately, large diff erences exist between the types and quality of data available for each country, and the extent to which the data can be made public also varies widely. some countries only have data available from traps for buoyant fl uids, where the tasr will be low not taking into account any potential for storage outside such traps by residual saturation. some countries have included aquifer formation data; here the tasr calculation will be more meaningful. in the great majority of countries, uncertainties related to lack of reservoir parameter data also remain. th e acquisition of such data will potentially require a sustained campaign of geological mapping and characterisation of storage capacity, or at least signifi cantly more time and fi nancial resources to assemble and enter all available data. th ese factors limit the results obtained from the co2stop project and it is recommended that further resources are made available for improving the results. in a european context, the technically accessible co2 storage resource (tasr) or theoretical storage resource should only be used for extra-european international resource comparisons because it is clear that the tasr is several times greater than the practical co2 storage capacity. consequently quoting the tasr can be misleading, giving false impressions of capacity if a critical distinction between resource and reserve estimates is not made. co2stop gis data analysis tool co2stop data interrogation system co2 injection capacity data entry system view, filter and export data import results dsf: deep saline formation dgf: depleted gas field fill data calculations & data choices primary input output data storefit areas studied for the co2stop gis st o ra ge id c ap ac it y d at ab as e (m t) fi el d e xe n t (k m 2 ) t h ic kn es s (m ) st o r a g e u n it t r a p t r a p n a m e c a se n a m e c a se n o messages input file output file countries studied countries not participating in co2stop project aquifer daughter units hydrocarbon daughter units storage units formations length name 62 89 fig. 2. schematic representation of the database analysis & interrogation tool, showing the gis and the storefit monte carlo analysis tool. arrows indicate data exchange between the separate elements of the tool. the map shows the reported resources in the co2stop project. 8888 conclusions th e calculations of co2 storage locations throughout europe made by the co2stop project database paint a broad picture, but also identify the gaps in our knowledge. th ese gaps must be fi lled with further data entry and, potentially, new geological studies, seismic surveys and drilling must be undertaken to make more precise data available. a common european legislation allowing equal access to proprietary subsurface information would be benefi cial for this purpose. it is critically important to understand the assumptions that lie behind the storage capacity estimates. th ese are especially relevant for saline formations, the capacities of which were derived without taking regulatory or economic limitations into account. th e co2stop method has made signifi cant progress towards establishing probabilistic estimates of the co2 storage resource in europe in a way that will allow comparisons with other regions of the world, and which will also be useful to policy makers. however, the partial data entry into the project database means that the current project only marks the beginning of the process of resource estimation and certainly not the end. acknowledgements th e co2stop project was funded by the european commission (project no ener/c1/154-2011-si2.611598). we express our sincere thanks to andrei bocin-dumitriu (ec joint research centre) and to kai tullius, øivind vessia, rakel hunstad and ilinca balan from the european commission, directorate general for energy for their help and support with this project and to this report and the other deliverables. we also thank the co2stop project partners for their contributions of country specifi c information. legal notice th is publication is based on a project for the european commission; however it refl ects the views only of the authors, and the commission cannot be held responsible for any use which may be made of the information contained therein. references bachu, s., bonijoly, d., bradshaw, j., burruss, r., christensen, n.p. holloway, s. & mathiassen, o.m. 2007: estimation of co2 storage capacity in geological media – phase 2. work under the auspices of the carbon sequestration leadership forum (www.cslforum.org). final report from the task force for review and identifi cation of standards for co2 storage capacity estimation, 43 pp. washington: carbon sequestration leadership forum. blondes, m.s. et al. 2013: national assessment of geologic carbon dioxide storage resources –methodology implementation. u.s. geological survey open-file report 2013–1055, 26 pp., http://pubs.usgs.gov/ of/2013/1055/ brennan, s.t. 2014: th e u.s. geological survey carbon dioxide storage effi ciency value methodology: results and observations. energy procedia 63, 5123–5129. brennan, s.t., burruss, r.c., merrill, m.d., freeman, p.a. & ruppert, l.f. 2010: a probabilistic assessment methodology for the evaluation of geologic carbon dioxide storage. u.s. geological survey open-file report 2010/1127, 31 pp. heidug, w. (ed.) 2013: methods to assess geologic co2 storage capacity: status and best practice, workshop report, 44 pp. paris: international energy agency. iea 2008: energy technology perspectives: scenarios and strategies to 2050, 650 pp. paris: international energy agency. iea 2013: technology roadmap. carbon capture and storage, 63 pp. paris: international energy agency. metz, b. et al. (eds) 2005: carbon dioxide capture and storage. ipcc 2005, 431 pp. cambridge university press. u.s. geological survey geologic carbon dioxide storage resources assessment team 2013: national assessment of geologic carbon dioxide storage resources – results (ver. 1.1, september 2013). u.s. geological survey circular 1386, 41 pp. van der meer, l.g.h. & van wees, j.d. 2006: eff ects of co2 solubility on the long-term fate of co2 sequestered in a saline aquifer. th e leading edge 25, 1276–1280. authors’ addresses n.p., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: nep@geus.dk a.b.-d., european commission, dg jrc, institute for energy and transport, energy technology policy outlook unit, westerduinweg 3, 1755 le petten, the netherlands. s.h. & k.k., british geological survey (bgs), kingsley dunham centre, keyworth, nottingham, ng12 5gg, uk. f.n., tno, earth environment and life sciences, postal address: p.o. box 80015, 3508 ta utrecht, the netherlands. n.s., british geological survey, murchison house, west mains road, edinburgh, eh9 3la, uk. geological survey of denmark and greenland bulletin 33, 2015, 17-20 17 thrust-fault architecture of glaciotectonic complexes in denmark stig a. schack pedersen and lars ole boldreel cross sections of glaciotectonic complexes are exposed in coastal cliff s in denmark, which allow structural studies of the architecture of thin-skinned thrust-fault deformation (pedersen 2014). however, the basal part of the thrust-fault complex is never exposed, because it is located 50 to 100 m below sea level. it is in the basal part the most important structure – the décollement zone – of the complex is found. th e décollement zone constitutes the more or less horizontal surface that separates undeformed bedrock from the displaced thrust-sheet units along the décollement level. one of the most famous exposures of glaciotectonic deformations in denmark is the møns klint glaciotectonic complex. th e structures above sea level are well documented, whereas the structures below sea level down to the décollement level are poorly known. modelling of deep structures was carried out by pedersen (2000) but still needs documentation. a glaciotectonic c omplex aff ecting comparable rock units, such as the chalk at møns klint, was recently recognised in seismic sections from jammerbugten in the north sea (fig. 1). th ese sections provide an excellent opportunity for comparable studies of the upper and lower structural levels in thin-skinned thrust-fault deformation, which is discussed in this paper with examples from three major glaciotectonic complexes. architecture of thrust-fault deformations in glaciotectonic complexes in contrast to fold-belt ranges, glaciotectonic complexes are relatively small and therefore easier to study. for these structural complexes, an architectural classifi cation was defi ned based on description and ordering of surfaces and their relations (pedersen 2014). it is emphasised that the creation of constructions comprising surfaces is the basic element of their architecture. in geology the use of architectural analysis is well known in investigations of sedimentary deposits. a concept for ordering of bounding surfaces in the architecture of aeolian dunes was suggested by brookfi eld (1977), and a similar concept was suggested by miall (1985) for facies analysis of fl uvial deposits. for the analysis of glacial architecture and construction of 3d geological models of glaciotectonic complexes the classifi cation of a hierarchy of bounding surfaces comprises four orders of surfaces (pedersen 2014). th e décollement surface is defi ned as a fi rst-order surface (fig. 2). th e décollement surface is the ‘base’ of the complex, and therefore the top of the complex also has to be defi ned as a fi rst-order surface. th is second fi rst-order surface is the topographic top of the tectonic complex, or alternatively, a truncating unconformity, above which post-deformational units occur. th e internal framework of a tectonic complex comprises thrust sheets. th ese are bounded by thrust faults, which are defi ned as second-order surfaces (fig. 2). th e thrust faults are diff erentiated into ramps and fl ats, where a ramp cross-cuts the bedding, whereas the fl at is more or less bedding-parallel. when two or more thrust sheets are bounded by ramps and fl ats they form duplexes. th ese generally form imbricate complexes or may be stacked so they form complex repetitions of the geological units (as exemplifi ed by pedersen 2005). th e folded beds comprise third-order surfaces. th ese are diff erentiated into anticlines, synclines, recumbent folds and monoclinal bends. folds may further be classifi ed from the © 2015 geus. geological survey of denmark and greenland bulletin 33, 17–20. open access: www.geus.dk/publications/bull weichselian maximum saalian maximum baltic sea north sea jb mk fk 300 km fig 1. extent of ice sheets during the two last glaciations and the location of the three glaciotectonic complexes mentioned in this paper. jb: the jammerbugt glaciotectonic complex was formed by an ice advance from central scandinavia during the saalian. fk: the fur knudeklint glaciotectonic complex was formed during an ice advance from norway during the late weichselian. mk: the møns klint glaciotectonic complex is exposed in a coastal cliff by the baltic sea, and it was formed during the latest part of the weichselian. 1818 orientation of their axial surface, the angle of their limbs and the inclination of their fold axes. fourth-order surfaces include all small-scale structures such as faults with small displacements; such faults are important for the understanding of the dynamic development. joints and anastomosing joints indicate early deformation impact, and the zone axis of conjugate faults indicates the direction of compaction. th e asymmetry of smalland mesoscale folds and the sense of displacement on faults as indicated by groove marks can be used to reconstruct the kinematics of deformation. for the macro-scale identifi cation of the head and tail of glaciotectonic complexes, a distal, a central and a proximal domain are defi ned. th e domain nearest to the foreland (the head end) is regarded as the distal part, which is commonly limited by the trace of the last thrust fault displaced towards the undeformed foreland (the tipline). th e central domain displays the bulk architecture of the complex. th e proximal domain comprises the deepest level of deformation with the most complicated structural relationships, potentially including superimposed deformation and mud diapirism. situated at the tail end, the proximal domain is delimited by the contact to the hinterland of the complex. for glaciotectonic complexes, the hinterland contact is the boundary between the hill and the hole in a ‘hill/hole pair’. at the time of dislocation, it formed the contact between the pushing ice and the dislocated geological units. the distal domain in the distal part of a glaciotectonic complex, the dip of the thrust fault ramps is gentle and the thrust fault fl ats are almost horizontal (fig. 3). th e thrust sheets are thinner than in the central domain, which is a consequence of the décollement surface that rises from the deepest level in the trailing end to the topographic surface in the foreland. one of the most surprising features of distal thrust sheets is their length. in the seismic section from the jammerbugt glaciotectonic complex the length of a thrust sheet exceeds 1 km, and the thrust sheets in the northern part of møns klint are more than 500 m long (fig. 3). such long thrust sheets are surprising when their thickness is taken into account. at møns klint the chalk sheets in the distal part of the glaciotectonic structure are only c. 25 m thick and one would expect that the forces pushing the thrust sheets would break them up into fragments. th e explanation for this missing fragmentation is that high porewater pressure along the thrust faults carries the unbroken thrust sheets. the central domain when a long and relatively thin thrust sheet is created in the distal part of a complex it is easy to understand that, when the thrusting propagates, the distal domain will move to the central domain during the formation of a new distal domain next to the foreland. two marked types of structures may form during this development: (1) the thrust sheets are broken into shorter segments creating imbricate fans along steeper-dipping thrust faults (pedersen 2005), or (2) super 3 1 3 2 1 1 1 2 3 3 2 2 2 1 sea level a b 50 m glacial deposits røsnæs clay formation stolleklint clay holmehus formation ash layers +1 to +13 ash layers –13 to +1 ash layers –33 to –13 fur formation north southc fig. 2. the fur knudeklint glaciotectonic complex with ash layers in the eocene diatomite of the fur formation that was deformed by the norwegian ice advance in the late weichselian. a: an anticline, a syncline and steeply dipping layers. b: imbricate duplexes. c: schematic section. 1: first order surfaces, the décollement surfaces at the base and the glaciotectonic unconformity at the top. 2: second-order surfaces, the thrust faults that divide the glaciotectonic complex into thrust-sheet segments. 3: thirdorder surfaces, the fold structures. to illustrate the typical third-order surfaces the hanging-wall anticlines have been extended above ground. fourth-order surfaces are too small to be illustrated, but are documented in pedersen (2014). the imbricate duplexes in the southern part of the complex are also seen in b. 19 posed thrust sheets are displaced together and passing over new, more deeply seated ramps. during this translation an antiformal stack is created, which is the explanation for the impressive structure at dronningestolen at møns klint (pedersen 2000, 2014; pedersen & gravesen 2009). a similar structure has been identifi ed in a seismic section from the central domain of the jammerbugt glaciotectonic complex. the proximal domain th e proximal part of a thin-skinned thrust-fault complex is characterised by an increasing number of thrust fault ramps and fl ats, imbricate thrust sheets and duplex segments (pedersen 2005). in fig. 4 this is illustrated by a section from the proximal part of the jammerbugt glaciotectonic complex and the southernmost imbricate thrust sheets at møns klint. th e thrust sheets at møns klint are c. 60 m thick, and the dips of the thrust faults are close to the maximum angle of fracturing (< 45º). th e thrusting probably includes superimposed tilting on deeper thrust faults below sea level. according to surlyk (1984) the stratigraphic level of the maastrichtian chalk is lower in the thrust sheets shown in fig. 4 than the chalk exposed in the distal domain in fig. 3. th us the thrusting and hence also the position of the décollement surface have shift ed to a deeper level in the proximal domain. th is relationship is also seen in the thrust-fault architecture of the seismic section from jammerbugten (fig. 4). in the distal and central domains the décollement surface is situated above the base of the chalk group (bc in fig. 4). in the proximal domain, the décollement surface drops down to the lower part of the marked refl ectors representing the base of the chalk group. th e marked bc refl ectors are present in the thrust sheets of the tailing part of the proximal domain. th e thrust sheets in the jammerbugt complex are about twice as thick as the thrust sheets at møns klint. th is refl ects that at møns klint only the frontal parts of the wedge-shaped thrust sheets are exposed, whereas in the seismic section the deeper, thicker parts of the thrust sheets can be recognised. conclusion th e architecture of thin-skinned thrust-fault deformation is described on the basis of three glaciotectonic complexes. th e thrust-fault architecture of thrust-fault belts and of glaciotectonic complexes is fairly similar even though the former are related to compressional regimes in plate-tectonic settings, the latter to compression caused by gravitational expansion of ice sheets. glaciotectonic thrust-fault complexes are divided into proximal (nearest to the source of force), 50 m b 500 m 200 m a bc fig. 3. thrust-fault architecture in the distal domains of two complexes. a: seismic section from the jammerbugt glaciotectonic complex. the strong reflectors are interpreted as the base of the chalk group (bc) in the north sea. this implies that the main parts of the thrust sheets comprise upper cretaceous chalk. b: thin, gently dipping thrust sheets in the northern part of møns klint. the chalk at møns klint is of maastrichtian age. 2020 central and distal domains (farthest away from the source of force). th e distal domain includes the foreland boundary of the thrust-fault complex, and it is characterised by long and thin, gently dipping thrust sheets. th e central domain is characterised by sequentially superimposed folding of thrust sheets formed in the distal domain. imbricate thrustfault segments are formed when the sheets break. th e proximal domain is characterised by the shift of the décollement surface down to the deepest level, thicker thrust sheets and stacking of thrust-fault duplexes. references brookfi eld, m.e. 1977: th e origin of bounding surfaces in ancient aeolian sandstones. sedimentology 24, 303–332. miall, a.d. 1985: architectural-element analysis: a new method of facies analysis applied to fl uvial deposits. earth-science reviews 22, 261–308. pedersen, s.a.s. 2000: superimposed deformation in glaciotectonics. bulletin of the geological society of denmark 46, 125–144. pedersen, s.a.s. 2005: structural analysis of the rubjerg knude glaciotectonic complex, vendsyssel, northern denmark. geological survey of denmark and greenland bulletin 8, 192 pp. pedersen, s.a.s. 2014: architecture of glaciotectonic complexes. geosciences 4, 269–296. pedersen, s.a.s. & gravesen, p. 2009: structural development of maglevandsfald: a key to understanding the glaciotectonic architecture of møns klint, se denmark. geological survey of denmark and greenland bulletin 17, 29–32. surlyk, f. 1984: th e maastrichtian stage in nw europe, and its brachiopod zonation. bulletin of the geological society of denmark 33, 217–223. author’s addresses s.a.s.p., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: sasp@geus.dk l.o.b., department of geosciences and natural resource management, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. 500 m bc 200 m 50 m a b fig. 4. two examples showing thrust-fault architecture in the proximal parts of glaciotectonic complexes. a: seismic section from the jammerbugt glaciotectonic complex. the décollement surface is stepping down to the lower part of the marked reflectors that represent the base of the chalk group (bc). b: the oldest chalk (surlyk 1984) at møns klint is found in the centre of the photograph, which shows the southernmost imbricate thrust sheets in the møns klint glaciotectonic complex. geological survey of denmark and greenland bulletin 26, 2012, 61-64 61© 2012 geus. geological survey of denmark and greenland bulletin 26, 61–64. open access: www.geus.dk/publications/bull geological assessment of the east greenland margin michael b.w. fyhn, thorkild m. rasmussen, trine dahl-jensen, willy l. weng, jørgen a. bojesen-koefoed and tove nielsen the east greenland margin consists of a number of sedimentary basins, platforms and structural highs (figs 1, 2). due to the challenges imposed by the arctic climate, the region is in an early stage of exploration, and knowledge of the geology and petroleum potential of the margin is limited. however, the significant prospectivity of the conjugated european north atlantic margin and the nature of the northeast greenland onshore geology prompt for future offshore exploration. the us geological survey thus highlighted the north-east greenland margin in their latest assessment of the arctic region (gautier et al. 2011). with a mean estimate of undiscovered recoverable oil, gas, and natural gas liquids of approximately 31 billion barrels of oil equivalents, the us geological survey ranked the north-east greenland margin fourth in the entire arctic region, only superseded by known producing petroleum provinces. in preparation for the initial east greenland licence rounds in 2012 and 2013 the geological survey of denmark and greenland gathers geological information on the margin necessary for the decision process of the greenland authorities regarding exploration. geophysical analyses complemented by well-data, onshore geology and information from the conjugated atlantic margin form the backbone of the study. the east greenland margin is covered by an open seismic grid supplemented by gravimetric and magnetic data. all existing 2d seismic, gravimetric and magnetic data are included in the current study. most of the data are confidential. restricted by the general confidential nature of the project, this paper aims to summarise the geology of the east greenland margin based on the current and previous studies and to briefly assess some of the implications for the regional petroleum prospectivity. east greenland offshore geology the east greenland margin stretches almost 3000 km from kap farvel in the south to nordostrundingen in the north (fig. 1). in the south it is narrow, but north of liverpool land it begins to widen to more than 300 km. the geology varies considerably along its length. east greenland margin south of shannon a continuous paleocene–eocene (61–53 ma) preto postbreak-up volcanic cover blankets almost the entire east greenland margin between kap farvel and the island of shannon (fig. 1). paleocene – early eocene magmatism refig. 1. the east greenland margin from kap farvel to nordostrundingen with offshore magnetic total field anomaly from the camp-gm compilation (gaina et al. 2009). -1 250 km nordostrundingen shannon liverpool land blosseville kyst fig. 3 odp 987e fig. 3 odp 987e odp 917aodp 917a kap farvel continent–ocean transition hold with hope –264 –156 –112 –83 –59 –37 –16 7 23 39 59 82 110 174 345 [nt] ffffffffffffffffffffffffffffffffffffffffffffff ppppppppppp 999999999pp 9999999p 9ppppppp 70°n 10°w 6262 cords the trace of the icelandic hotspot that induced magmatism along most of the proto-north atlantic margin and likely led to continental break-up between europe and greenland. thick seaward-dipping reflection packages interpreted as basalt flows exist near the continent–ocean transition indicating the presence of a very thick basalt-dominated succession below the distal part of the margin (fig. 3). basalt successions several hundred metres to kilometres thick exist onshore central east and north-east greenland. it is likely that similar and thicker paleocene – lower eocene basalt successions cover large parts of the offshore margin judging from seismic data, wells and onshore analogies (fig. 2). only the odp well 917a offshore south-east greenland penetrates the entire basaltic cover (fig. 1; larsen & saunders 1998). the well was drilled on the shoreward part of the margin and encountered 779.5 m of mainly basalts and thus documents the presence of a thick volcanic cover in the shoreward part of the margin that, according to seismic evidence, significantly thickens seawards. offshore liverpool land, the paleocene/eocene top-basalt surface is not clearly discernable on all seismic transects. however, this is probably due to the deep burial of the basalt succession beneath a thick post-basaltic cover. in the area offshore hold with hope, the presence of volcanics younger than the regional upper paleocene – lower eocene flood basalts is interpreted from seismic data (fig. 2). these are likely to be genetically related to middle eocene – earliest miocene magmatism observed immediately onshore. sub-basalt sediments. the tertiary basalt cover generally impedes seismic imaging of the deeper geology. offshore southeast greenland, sub-basaltic acoustic velocities derived from refraction seismic data generally range from 4.5 to 6 km/s increasing downwards (hopper et al. 2003), which does not provide unequivocal evidence for regional, thick sub-basaltic sedimentary basins. a sub-basaltic succession of deformed metasediments barren of fossils drilled in odp well 917a was suggested to be upper cretaceous to lower paleocene in age, like onshore sediments farther north (larsen & saunders 1998). however, the succession could also represent much older caledonian metasediments formed in a foreland setting, comparable to onshore metasediments farther to the north and lying beneath the conjugated british margin. offshore north-east greenland, upper palaeozoic – mesozoic deposits below the basalts are suggested by the presence of thick such deposits along the shore, immediately onshore and on the conjugated norwegian north atlantic margin. moreover, a thick pre-basaltic sedimentary succession can be recognised offshore shannon, continuing southwards below the basalts. post-paleocene sediments. the eocene and younger succession between kap farvel and shannon attains thicknesses up to c. 3 sec two-way travel time (twt). the succession is absent or very thin shoreward, and over large areas, break-up volcanics and upper palaeozoic and caledonian basement crop out at the seafloor. farther offshore, an eocene and younger shelf progradation along most of the margin resulted in the build-up of the prominent modern shelf-slope that generally straddles the continent–ocean transition. the blosseville kyst fault zone along the blosseville kyst and its northern continuation the liverpool land fault zone downfault the break-up volcanic cover with up to 2 sec twt (fig. 2). rifting along blosseville kyst and livepool land fault zones was likely 200 km 25°w 15°w 15°w 5°w 70°n 75°n ? ? ? offshore ne greenland with licence area area covered by u. paleocene – l. eocene flood basalts u. paleocene – l. eocene basalts likely concealed underneath younger geology post l. eocene basalts area with u. paleocene – l. eocene intrusions and sporadic volcanics approximate continent– ocean transition zone ocean crust major normal fault anticlinal trend seismic line core hole 50 000 km2 licence area and blocks shannon t h e ti s b as ind a n m a r k sh a v n b a s i n k o ld e w e y p la t f o r m d an m ar k sh av n r id ge fig. 2. the extent of massive volcanism along the north-east greenland margin. magmatism influenced areas up to 200 km north of the basalt cover as indicated by the heavily intruded part of the thetis and danmarkshavn basins. 63 associated with a westward jump of the spreading axis and the separation of the jan mayen microcontinent from central east greenland during mid-cenozoic time (mjelde et al. 2008). faulting mostly affected the lower part of the postbreak-up succession, and a significant syn-tectonic fan delta flanks the fault zones suggesting (?)late eocene/oligocene to c. early miocene faulting (fig. 3). faults only penetrate the seafloor near the coast where erosion and non-deposition have prevented fault burial. east greenland margin north of shannon north of the island of shannon, the landward edge of the basalt cover curves to the east and continues farther seawards to the north where a wide basalt belt fringes the continent–ocean transition (fig. 2). this leaves the shelf area north of shannon virtually unaffected by tertiary flood basalts and thus permits fair to good seismic imaging of paleocene and older successions. as a result, up to c. 10 sec twt several deep sediment-filled basins can be recognised north of shannon that are likely to correspond to between 15 and 20 km of sedimentary infill. the area can be structurally divided into the danmarkshavn and thetis basins separated by the danmarkshavn ridge – a c. 300 km long rifted basement high (fig. 2). the danmarkshavn basin is downfaulted relative to the koldewey platform to the west and the danmarkshavn ridge to the east. the thetis basin marks the oceanward part of the north-east greenland margin that borders the north atlantic oceanic crust. the depositional succession filling in the thetis basin may be of comparative thickness to that of the danmarkshavn basin, but there are significant differences in the geological development of the two basins. the koldewey platform makes up an area of sedimentcovered, more shallow basement bridging the danmarkshavn basin and the onshore basement province. the danmarkshavn ridge stands in marked contrast to the deeply downfaulted thetis basin farther east. as the high degrades into fault blocks to the north, the ridge plunges and the structural separation of the thetis and danmarkshavn basins crumbles. structural and stratigraphic development. the deepest deposits in the danmarkshavn basin and on the danmarkshavn ridge form a prominent syn-rift succession interpreted to be devonian and carboniferous in age (hamann et al. 2005). the unit is located in a rift system often confined by lowangle extensional faults, which may reflect extensional reactivation of caledonian thrust faults. the syn-rift succession is overlain by a unit in the central and southern part of the danmarkshavn basin that has in places been mobilised into diapiric and pillow-like structures interpreted as salt kinetic features. the succession is interpreted as lower carboniferous to early permian deposits comparable to coeval sediments in the western part of the barents sea (hamann et al. 2005). the late palaeozoic successions are buried beneath an up to a few kilometres thick unit only moderately affected by syn-depositional faulting and interpreted as upper permian – jurassic deposits. by analogy to the conjugated atlantic margin and onshore geology, the upper permian – jurassic succession is likely to contain a number of prominent source rock intervals. of these the upper jurassic equivalent to the kimmeridge clay formation may be of particular importance to potential petroleum systems in the region. the top of the jurassic is marked by a pronounced angular unconformity associated with the onset of cretaceous rifting fig. 3. an e–w seismic profile crossing the liverpool land fault zone and related syn-tectonic fan delta deposits and the continent–ocean transition indicated by seaward-dipping reflectors (sdrs). the odp 987e well drilled recent to miocene strata and terminated slightly above the basalts. 10 kmtop of oligocene–miocene basalts 0 1 2 3 tw o -w ay t ra ve l ti m e (s ec ) 4 5 ew fan delta liverpool land fault zone sdrs fan delta liverpool land fault zone sdrs odp 987e 6464 (hamann et al. 2005). cretaceous rifting resulted in several kilometres of downfaulting towards the thetis basin, and to a lesser extent also along the margins of the danmarkshavn basin. consequently, the danmarkshavn ridge became an elevated rift-shoulder subject to erosion during most of the cretaceous. in contrast, vast amounts of sediment were shed into the adjacent basins and kilometre-thick cretaceous successions exist in the depocentres. large faultand anticlinal structures within cretaceous and older successions formed in association with the cretaceous extension. many of these resemble structures hosting world-class oil and gas fields on the conjugate european atlantic margin, thus boosting expectations for the north-east greenland margin. a change from fault-controlled subsidence concentrated along the basins to regional eastward tilting took place close to the cretaceous/tertiary boundary. the change resulted in flooding and burial of the danmarkshavn ridge during the early tertiary and instigated regional eastwards progradation across the basin areas. only modest volcanism, if any, took place across the basins north of shannon. however, the southern danmarkshavn and thetis basins were heavily intruded far north of the edge of the east greenland basalt province, which degrades seismic imaging of the deepest stratigraphy over large areas. during mid-tertiary time, sediment supply increased resulting in build-up and eastward progradation of a steep shelf-slope. this was likely caused by denudation of the east greenland margin and the inner basin areas. as a result, the lower tertiary and mesozoic deposits lie beneath the plio/ pleistocene sediments along the inner half of the north-east greenland shelf. implications for the petroleum prospectivity offshore east greenland assessment of the petroleum prospectivity along the southeast and central east greenland margin is associated with considerable uncertainties. a sub-basaltic source rock seems to be a prerequisite for a working petroleum system along the margin, but the thick volcanic cover complicates analysis of the deeper geology. regional presence of sub-basaltic basins is at present uncertain offshore south-east greenland, which also makes the prediction of an effective petroleum system along this part of the margin uncertain. farther north between liverpool land and shannon the chances of sub-basaltic basins containing significant source-rock intervals are high. however, interpretation of the sub-basaltic structural style is problematic at present and the maturation history of potential source-rock intervals is uncertain, which complicates the prospectivity assesment. farther north, extreme sea-ice conditions are a major impediment to future exploration. in addition, tertiary uplift and erosion as well as past glacial advances across the shelf may have affected the regional petroleum potential. however, the presence of large structures and the good chance of adequately matured source-rock intervals urge for further investigation. the highest density of structures is located along the flanks of the basins and across the danmarkshavn ridge, which are also the areas covered in the coming licence rounds (fig. 2). many of these structures are associated with significant potential direct hydrocarbon indicators, which may reflect the presence of significant petroleum accumulations offshore north-east greenland. this is highly encouraging for further exploration in one of the last frontier areas on earth. acknowledgement the bureau of minerals and petroleum, nuuk, is acknowledged for funding. references gaina, c., werner, s.c. & the camp-gm group 2009: circum-arctic mapping project – gravity and magnetic maps. report 2009.010, 21 pp. trondheim: geological survey of norway. gautier, d.l. et al. 2011: assessment of ne greenland: prototype for development of circum-arctic resource appraisal methodology. in: spencer, a.m. et al. (eds): arctic petroleum geology. geological society memoirs (london) 35, 663–672. hamann, n.e., whittaker, r.c. & stemmerik, l. 2005: geological development of the northeast greenland shelf. in: doré, a.g. & vining, b.a. (eds): petroleum geology: north-west europe and global perspectives. proceedings of the 6th petroleum geology conference, 887–902. london: geological society. hopper, j.r., dahl-jensen, t., holbrook, w.s., larsen, h.c., lizarralde, d., korenaga, j., kent, g.m. & kelemen, p.b. 2003: structure of the se greenland margin from seismic reflection and refraction data: implications for nascent spreading center subsidence and asymmetric crustal accretion during north atlantic opening. journal of geophysical research 108, 2269, http://dx.doi.org/10.1029/2002jb001996 larsen, h.c. & saunders, a.d. 1998: tectonism and volcanism at the southeast greenland rifted margin: a record of plume impact and later continental rupture. in: saunders, a.d., larsen, h.c. & wise, s.w. (eds): proceedings of the ocean drilling program scientific results 152, 503–533. mjelde, r., raum, t., breivik, a.j. & faleide, j.i. 2008: crustal transect across the north atlantic. marine geophysical research 29, 73–87. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: mbwf@geus.dk geological survey of denmark and greenland bulletin 1, 217-230 217 the jurassic of the netherlands g.f. waldemar herngreen, wim f.p. kouwe and theo e.wong a recent revision of the lithostratigraphy of the netherlands has triggered an extensive re-evaluation of existing ideas on the jurassic structural and depositional history. significant advances can be attributed to the incorporation of sequence stratigraphic concepts. in the course of the triassic and jurassic, structural complexity increased progressively. the jurassic sedimentary succession can be subdivided into three depositional megasequences. megasequence i (rhaetian– aalenian) reflects the period between the so-called early and mid-cimmerian tectonic phases. megasequence ii (aalenian – middle callovian) covers the period of activity of the mid-cimmerian phase. megasequence iii (middle callovian – ryazanian) corresponds with the period between the mid-cimmerian and late cimmerian phases (particularly after pulse ii). in this latter megasequence, six stages (iiia–f) are recognised. sediments deposited during the rhaetian and ryazanian bear a stronger affinity with the jurassic succession than with triassic and cretaceous sediments respectively. these stages are thus treated here as an integral part of the jurassic succession. during the rhaetian–bajocian the area subsided relatively uniformly. a sheet of predominantly fine-grained marine sediments of great lateral uniformity was deposited. during the toarcian, in particular, basin circulation was largely restricted. the cooling that followed the thermal central north sea dome uplift triggered an important extensional phase during the aalenian–callovian. the rift phase resulted in the formation of several smaller basins, each with its own characteristic depositional succession. the basins fall into three structural provinces: the eastern province (lower saxony basin, e–w-striking); the northern province (central graben, n–s-striking); and the southern–central system (roer valley graben – broad fourteens, with a strong nw–se strike). the mid-cimmerian event started to affect the dutch basins during the bajocian. sedimentation ceased in the dutch central graben while it persisted in a predominantly coarse-grained, shallow marine facies in the southern basins (roer valley graben, west netherlands basin). extensional tectonics in the central graben were initiated during the middle callovian, with the deposition of continental sediments. during the oxfordian–kimmeridgian, marine incursions gradually became more frequent. marine deposition in the other basins in the south persisted into the oxfordian, at which time deposition became predominantly continental. marine conditions gradually returned in the south during the ryazanian–barremian, with a series of advancing partial transgressions from the north. the present-day distribution of jurassic strata in the netherlands was determined largely by erosion associated with late cretaceous – paleocene uplift. keywords: the netherlands onshore and offshore, jurassic, lithostratigraphy, sequence stratigraphy, tectonics, regional geology g.f.w.h.* & t.e.w., netherlands institute of applied geoscience tno – national geological survey, p.o. box 80015, nl-3508 ta utrecht, the netherlands. *retired. e-mail: g.f.w.herngreen@bio.uu.nl w.f.p.k., wintershall noordzee b.v., eisenhowerlaan 146, nl-2517 jl the hague, the netherlands. geological survey of denmark and greenland bulletin 1, 217–229 (2003) © geus, 2003 the jurassic succession in the netherlands has significant hydrocarbon source rock and reservoir potential and the stratigraphy of the succession is thus a subject of particular interest. the jurassic lithostratigraphic nomenclature in the netherlands has recently been updated in a joint effort by the geological survey of the netherlands (rgd) and nogepa, the organisation of oil companies active in the netherlands. this project has benefited significantly from integration of new knowledge on jurassic biostratigraphy. in addition, sequence stratigraphic concepts have been used to explain the observed distribution of facies in time and space. the results have been published in part in van adrichem boogaert & kouwe (1994–1997). in this study, we treat the most recent ideas about the jurassic geological history of the netherlands. the geological history of the jurassic in the netherlands has been presented previously by haanstra (1963), heybroek (1974) and van wijhe (1987). ziegler (1990) treated the area from a northwest european perspective, whereas burgers & mulder (1991) summarised certain aspects of the late jurassic and cretaceous history. regional stratigraphic overviews of the jurassic in the southern north sea have been presented by brown (1990) and cameron et al. (1992). michelsen & wong (1991) presented a stratigraphic correlation between the dutch, danish and southern norwegian sectors of the central graben. underhill & partington (1993) discussed the effect of middle jurassic thermal doming on regional jurassic stratigraphy in a sequence stratigraphic context. kimmeridgian to ryazanian sequence stratigraphy, biostratigraphy and the distribution of the main reservoir intervals in the north sea were dealt with by partington et al. (1993). during the last two decades, several authors from the geological survey of the netherlands (rgd) have published detailed review papers on the middle–late jurassic history of the dutch central (north sea) graben, the vlieland basin and adjoining onshore areas: herngreen & de boer (1984), herngreen & wong (1989), herngreen et al. (1988, 1991), wong et al. (1989), geological atlas of the subsurface of the netherlands, sheets i–v (rgd 1991a, b, 1993a, b, 1995) and x (nitg–tno 1998; map sheet vi is in press and sheets vii and viii are currently in preparation). the accumulated rgd knowledge, combined with contributions by several oil companies, served as a basis for the revision of the existing lithostratigraphic nomenclature for the netherlands (van adrichem boogaert & kouwe 1994–1997). the revision of the jurassic involved two working groups: permian – middle jurassic and upper jurassic – lower cretaceous. a correlation of the current dutch jurassic – lower cretaceous lithostratigraphy with the stratigraphic subdivisions of the neighbouring countries is presented in figure 1. structural setting during triassic–jurassic times, the structural style of the netherlands gradually changed from the single, extensive southern permian basin into a complex of smaller, largely fault-controlled, highs and lows. this transformation, related to the break-up of pangea (ziegler 1990), took place in several discrete extensional phases: scythian (hardegsen), norian–rhaetian (early cimmerian), aalenian – callovian/oxfordian (mid-cimmerian) and kimmeridgian– valanginian (late cimmerian). the intervening periods were dominated by regular thermal subsidence. the result can be seen by comparing figures 2 and 3, which display the triassic and late jurassic basin configurations, respectively. in the northern half of the netherlands’ offshore territory, the segregation between highs and basins was enhanced by halokinesis in the permian–triassic successions along the structural boundaries. associated salt withdrawal commonly determined the distribution of jurassic depocentres. throughout the jurassic, the area comprised three interactive structural provinces (fig. 3). 1. the lower saxony basin system (previous ems low), extending into central germany. 2. the north sea central graben – vlieland basin system, trending north–south across the mid-north sea high – ringkøbing–fyn high complex. 3. a block-faulted basin system extending nw–se through the netherlands, connecting the roer valley graben, the central netherlands basin, the west netherlands basin and the broad fourteens basin (previous off-holland low), and linked with the sole pit basin in the uk offshore. during the early jurassic, the area subsided relatively uniformly. a sheet of open marine, fine-grained sediments of great lateral uniformity was deposited. structural complexity gradually increased during early and middle jurassic times, and reached a maximum in the callovian. during the middle and late jurassic, in particular, each of the above-mentioned provinces accumulated its own characteristic depositional succession. the impact of the uplift of the central north sea thermal dome during the aalenian–bathonian is most 218 219 w ie he ng eb .qu ar zi t po rt asa nd st ei n (m ac ro ce ph . s st ) r ed c ha lk f m sp ee to n c la y fm sp ils by sa nd st on e fm k im m er id ge c la y fm c or al lia n fm o xf or d c la y fm w in te rt on f m fl am m en m er ge l cromer knoll group schieland group delfland subgroup rijnland group cromer knoll group scruff group central graben subgroup rijnland group osningsandstein niedersachsen gp breeveertien fm zurich fm humber group west sole group altena group altena group central graben gp keuperliasdoggermalm lias group penarth group h ol la nd f m h ol la nd f m v lie la nd c la ys to ne f m v lie la nd sa nd st on e fm v lie la nd sa nd st on e fm sc ru ff g re en sa nd f m k im m er id ge c la y fm fr ie se fr on t fm sl ee n fm a al bu rg f m po si do ni a sh al e fm po si do ni a sh al e fm po si do ni en sc hi ef er w er ke nd am f m w er ke nd am f m br ab an t fm r ød by f m so la f m t ux en f m v al ha ll fm fa rs un d fm lo la f m fj er ri ts le v fm w in te rt on f m sl ee n fm a al bu rg f m m id dl e g ra be n fm lo w er g ra be n fm r hä tk eu pe r v yl fm po ul fm h en o fm br yn e fm lo w er g ra be n sa nd fm w ei te ve en f m c oe vo rd en f m v lie la nd c la ys to ne fm m id dl e g ra be n sh al e fm bü ck eb er g fm ( ‘w ea ld en ’) n eo ko m -s ch ie fe rt on (h ils to n fm ) m in um us to n be nt he im er sa nd st ei n g ild eh äu se r sa nd st ei n r ot he nb er gsa nd st ei n se rp ul it m b m ün de r fm ei m be ck hä us er p la tt en ka lk g ig as sc hi ch te n k im m er id ge ka lk e & m er ge l h ee rs um er s ch ic ht en o rn at en to n m ac ro ce ph al en to n o sn in g gr øn sa nd pu zz le h ol e fm u pp er g ra be n fm a sp id oi de ssa nd st ei n w ür tt em be rg ic us sa nd st ei n ‘c or nb ra sh ’ a ge a lb ia n a pt ia n ba rr em ia n h au te ri vi an v al an gi ni an r ya za ni an po rt la nd ia n k im m er id gi an (s en su la to ) o xf or di an c al lo vi an ba th on ia n ba jo ci an a al en ia n t oa rc ia n pl ie ns ba ch ia n si ne m ur ia n h et ta ng ia n r ha et ia n u ni te d k in gd om so ut he rn n or th s ea so ur ce : r hy s (1 97 4) , c am er on e t a l. (1 99 2) d en m ar k d an is h c en tr al g ra be n so ur ce : j en se n et a l. (1 98 6) , m ic he ls en & w on g (1 99 1) n et he rl an ds co m po si te fo r th e no rt he rn , ce nt ra l a nd w es te rn s ec to rs n et he rl an ds (n or th -) ea st er n se ct or g er m an y lo w er s ax on y ba si n (w es te rn p ar t) so ur ce : b ri nk m an n (1 95 9) , c as ey e t a l. (1 97 5) , k em pe r( 19 76 ), k la ss en ( 19 84 ) epoch early cretaceous early jurassic l. triassicmiddle jurassiclate jurassic k or al le noo lit h nieuwerkerk fm ? ? = h ia tu s fi g. 1 . r eg io n al li th o st ra tig ra p h ic c o rr el at io n c h ar t o f t h e ju ra ss ic – l o w er c re ta ce o u s fo r th e n et h er la n d s an d n ei gh b o u ri n g co u n tr ie s. k im m er id gi an s en su la to eq u al s k im m er id gi an se n su a n gl ic o. f o r re ce n t re vi si o n s o f th e d an is h s ch em e, s ee m ic h el se n e t a l. (2 00 3, t h is v o lu m e) . m o d if ie d f ro m v an a d ri ch em b o o ga er t & k o u w e (1 99 4– 19 97 ). 220 evident in the northern offshore area of the netherlands, where lower jurassic successions are deeply truncated. the combination of initial cooling of the thinned crust under the dome and an overall transtensional tectonic regime triggered callovian–kimmeridgian rifting, resulting in the formation of the north sea central graben rift system. the uplift phase terminated the uniform subsidence and sheet-like deposition that characterised the early and middle jurassic. rifting induced structural differentiation into rapidly subsiding basins with high sediment infill rates and more stable, sediment-starved, platform areas. this differentiation persisted until the early cretaceous. outside the main basins depicted in figure 3, upper jurassic deposits are rare and thin, and associated with salt domes (rim synclines) or transverse fault zones. the north-eastern part of the netherlands formed the western fringe of the german ems low – lower saxony basin system. the ems low, a north–south elongated, 100 km sr o q a b e f g k 1 4 7 10 13 16 2 5 8 11 14 17 3 6 9 12 15 18 l m n p 3°e 5°e 7°e 54°n 52°n mid north sea high elbow spit h igh ringkøbing–fyn high c en tr al g ra be n sole pit basin em s lo w netherlands swell east netherlands high roer valley graben w est netherlands basin south hewett shelf london–brabant massif r he ni sh m as si f off-holland low structural high/subaerial landmass platform/occasionally flooded basin major normal fault (zone) seismic line shown in fig. 6 uk n dk g nl fig. 2. main triassic to early jurassic structural elements in the netherlands. the k-quadrant shows the block numbering system used for the netherlands’ continental shelf. each complete quadrant (1°n x 1°e) is divided into 18 blocks measuring 10´n x 20´e, numbered from nw to se. incomplete quadrants are numbered as if they were complete (e.g. the o-quadrant only comprises blocks 12, 15, 17 and 18). national sectors of the north sea: dk, denmark; g, germany; n, norway; nl, the netherlands; uk, united kingdom. modified from van adrichem boogaert & kouwe (1994–1997). fig. 3. middle–late jurassic to early cretaceous structural elements in the netherlands. for legend, see fig. 2. modified from van adrichem boogaert & kouwe (1994–1997). sr o q a b e f g k l m n p 52°n 54°n c en tr al g ra be n terschelling basin st ep g ra be n ijmuiden high gouwzee trough w est netherlands basin lower saxony basin hantum fault zone lauwerszee trough groningen high ameland block rifgronden fault zone schill grund high roer valley graben zandvoort ridge noord–holland platform texel–ijsselmeer high elbow spit h igh krefeld h igh maas bommel high peel block indefatigable fault zone gronau fault zone outer rough basin ringkøbing–fyn high mid north sea high cleaver bank high sole pit basin broad fourteens basin vlieland basin friesland platform winterton high central netherlands basin london–brabant massif rhenish massif vlieland high 3°e 5°e 7°e 100 km fault-bounded basin during carboniferous–triassic times, evolved into the slightly east–west-trending, sag-like lower saxony basin during the middle jurassic. a relatively undisturbed, continuous jurassic succession is found in the central parts of the basin in germany (fig. 1). in the eastern netherlands’ development, a hiatus spanning the aalenian – early kimmeridgian reflects the situation along the western margin of the basin. the west netherlands basin, the roer valley graben and the broad fourteens basin formed a different depositional province, connected with the sole pit basin of the uk sector of the north sea. since permian–triassic salts are largely absent in the subsurface here, these basins demonstrate a pure block-faulted nature. the three mentioned provinces merge in the onshore netherlands, where the central netherlands basin and the vlieland basin (a small pull-apart basin) were positioned. the transitions between these provinces are now obscured as a result of later erosion which has stripped jurassic sediments from most of the country. application of sequence stratigraphy in connection with the revision of the existing lithostratigraphic nomenclature for the netherlands (van adrichem boogaert & kouwe 1994–1997), sequence stratigraphic analysis was seen as a valuable additional tool for a better understanding of the complex geological settings and stratigraphic relationships. the lower and middle jurassic successions in the netherlands are of open marine origin, and developed in regionally uniform facies. depositional cycles are commonly discerned in this interval, and some of these were already incorporated in the old lithostratigraphic subdivision by nam & rgd (1980). in the recent revisions, therefore, it was not considered necessary to improve this scheme by incorporation of results from new biostratigraphic or sequence stratigraphic analyses. in the (middle–)upper jurassic, however, the common intercalation of continental and marine deposits makes sequence stratigraphy especially effective. therefore all available data (well and seismic correlations, biostratigraphic and palynological data) were integrated into a regional sequence stratigraphic framework. this framework was used to select those units which were suitable candidates for lithostratigraphic units. criteria for unit selection included mappability, relevance for exploration geology and the existence of diagnostic lithological and/or biostratigraphical criteria for their recognition. with the available data and time, the construction of a specific, local sequence framework fell outside the scope of the working group. instead, the group tied the recognised transgressive and regressive events to the best-fitting sequences in the cycle chart published by haq et al. (1988). this haq et al. (1988) time-scale has been recalibrated numerically to the scale of harland et al. (1990), which was chosen as the current standard for dutch stratigraphy. for the sake of completeness, figures 4 and 5 (facing pages 226 and 227, respectively) indicate the haq et al. (1988) sequence subdivision for the whole of the jurassic; figure 5 provides an overview of some significant biostratigraphic marker horizons that were used for calibration of dutch lithostratigraphy with the cycle chart. sequence stratigraphic analysis was carried out using concepts defined by the exxon school, viz. sequences are stratal units bounded by unconformities and their correlative conformities (van wagoner et al. 1988). in contrast, concepts of genetic sequence stratigraphy (sensu galloway 1989), in which genetic sequences are bounded by maximum flooding surfaces, are commonly invoked for regional correlation studies. an example of this approach is found in partington et al. (1993), for their regional correlation of the jurassic of the north sea area. it should be noted that their interpretation of dutch jurassic sediments was based on obsolete ideas derived from nam & rgd (1980). figure 4 presents the most modern ideas on the dutch chrono-lithostratigraphic development in the context of the development in the surrounding countries. depositional history the jurassic succession of the netherlands can be subdivided into three depositional megasequences. megasequence i (rhaetian–aalenian) reflects the period between the early and the mid-cimmerian tectonic phases. megasequence ii (aalenian – middle callovian) covers the period of activity of the mid-cimmerian phase. megasequence iii (middle callovian – ryazanian) corresponds to the period between the mid-cimmerian and the late cimmerian (pulse ii) phases (rgd 1991a). rhaetian–aalenian (megasequence i) this period was characterised by the deposition of a very uniform blanket of marine shales across large parts of northwest europe. in the netherlands, these deposits 221 are placed in the altena group. after a long period of restricted marine to continental deposition in the triassic, the last pulse of the early cimmerian extensional phase in the earliest rhaetian caused a marine transgression across large parts of europe. very fine-grained deposits of rhaetian age, containing abundant lacustrine and marine fossils, are placed in the sleen formation. hettangian–pliensbachian or lowermost toarcian sediments, consisting of a uniform succession of dark grey or black silty claystones with abundant pyritised fossil remains, are called the aalburg formation. towards the london–brabant massif (the southern basin fringe during the early jurassic), an increasing number of thin limestone beds are intercalated in this unit. basin circulation became restricted during the toarcian, with anoxic bottom waters and deposition of a bituminous shaly claystone, called the posidonia formation, over large parts of northwest europe. it constitutes the most prominent oil source rock in the netherlands. basin circulation returned to normal during the late toarcian – bajocian, as indicated by the deposition of the werkendam formation, consisting of marine silty mudstones and greensands. aalenian – middle callovian (megasequence ii) this was a period of significant tectonic activity, which caused important structural differentiation (figs 2, 3). the main effect of the tectonic phase traditionally addressed in the netherlands as ‘mid-cimmerian’ was the uplift of the central north sea dome. the deformation front gradually shifted southwards with time, resulting in a considerable delay between the timing of deformation of the dome crest in the central north sea (mid-aalenian) and the southernmost dutch provinces (kimmeridgian). the amount of intra-jurassic truncation decreases away from the dome (underhill & partington 1993), and consequently is most severe within the area of this study in the dutch central graben, the northernmost province of the netherlands, where open marine deposition of the altena group ceased in the bajocian. the dutch part of the central graben remained nondepositional during the bathonian – early callovian. in the middle callovian, deposition gradually resumed, in a continental facies. the uplift event probably also affected the terschelling basin, the dutch lower saxony basin and the central netherlands basin. however, the evidence for the impact of this phase was removed by later, more severe truncation in the jurassic successions of these basins. the impact of the mid-aalenian central north sea uplift is negligible in the southern basins (broad fourteens basin, west netherlands basin, roer valley graben). werkendam formation deposition continued, with an important bajocian influx of marine sandstones along the southern basin margin. depositional facies changed to shallow marine, sandy carbonates and marls during the bathonian (brabant formation, ‘cornbrash facies’). at least three carbonate–marl cycles were deposited in the bathonian to oxfordian, which are at present only preserved as erosional remnants in the southern basins. in the dutch achterhoek (eastern onshore), on the south-western fringe of the lower saxony basin, a special development is found (herngreen et al. 1984). the sedimentary succession up to the middle bathonian shows strong parallels with successions of the roer valley graben and the german lower saxony basin proper (fig. 1). the whole area appears to have been a single depositional province during this period. at the transition to the late bathonian, however, a differentiation occurred. during the late bathonian – callovian, an open marine claystone facies was deposited in germany and the achterhoek (informally termed the klomps member, fig. 4). in contrast, deposition of shallow marine sandy carbonates and marls in the roer valley graben and the west netherlands basin persisted into the oxfordian. middle callovian – ryazanian (megasequence iii) upper jurassic – lower cretaceous deposits in northwest europe demonstrate a step-by-step overall transgression from north to south. periods of marine transgression alternated with phases of short-term progradation of, partly tectonically controlled, continental siliciclastics. the continuous shifting of marine and terrestrial realms in response to syndepositional tectonism and sea-level fluctuations is reflected by the recurrent intertonguing of marine and continental sediments. the predominantly marine scruff group (upper jurassic – ryazanian) and rijnland group (lower cretaceous) interfinger with the mainly continental schieland group (upper jurassic – barremian) and the paralic to restricted marine niedersachsen group (upper jurassic – ryazanian). differential movement of fault blocks was caused by a combination of oblique-slip effects in an overall extensional regime and halokinesis in areas with salts in the subsurface (wong et al. 1989). together with a high sed222 223 iment input, this resulted in complex sediment distribution patterns and continuously shifting depocentres. the cyclic alternation of marine and continental sediments indicates that sediment deposition was also strongly controlled by sea-level changes. middle callovian – oxfordian (stage iiia) in the middle–late callovian, predominantly continental sedimentation (schieland group) resumed along the axis of the northern dutch central graben (fig. 4). the basal alluvial plain deposits are referred to the lower graben formation. this unit displays huge thickness variations due to onlap onto syndepositional topography, and due to differential subsidence. the depositional area gradually extended into the southern dutch central graben in the course of the latest callovian, oxfordian and kimmeridgian. deposition in the latest callovian started with the marginal marine rifgronden member of the friese front formation. it indicates connections to open marine areas in the northern central graben (central north sea), or possible links with the marine realm in the southern and central dutch basins. the shift to deposition of the middle graben formation in the latest callovian – early oxfordian reflects an abrupt change to more fine-grained, lakeand swamp-dominated sedimentation. this shift was accompanied by a short-lived, marine incursion into the whole dutch central graben, which is correlated with the maximum flooding surface of sequence lza4.1 of haq et al. (1988). during the middle oxfordian, lacustrine conditions prevailed throughout the dutch central graben. marine incursions were scarce and restricted to the far north during the rest of the oxfordian. the oxfordian was a period of significant transgression and onlap. areas adjacent to the northern dutch central graben, such as the southern dutch central graben, terschelling basin and step graben were incorporated into the depositional area. simultaneously, the open marine realm started to expand southwards into the dutch central graben, first into block f03 (middle–late oxfordian, sequence lza-4.2 of haq et al. 1988). initially, a stacked prograding coastal-barrier sand complex (upper graben formation) developed, behind which a paralic delta plain formed, represented by the deposits of the puzzle hole formation (blocks f08–f14). north of the barrier complex, the kimmeridge clay formation represents the open marine environment. the coastal sand bodies of the upper graben formation in the area of block f03 drowned at the end of the oxfordian (sequence lza-4.4 of haq et al. 1988). by this time, kimmeridge clay deposition also invaded the step graben, onlapping the upper permian zechstein group. non-marine deposition of the schieland group had already commenced in the dutch central graben during the callovian, but deposition of the marine brabant formation (altena group, oisterwijk limestone member) persisted into the early–middle oxfordian in the southern netherlands (haanstra 1963; nam & rgd 1980). cessation of marine altena group deposition in the roer valley graben, the west netherlands basin, the broad fourteens basin and probably the central netherlands basin – lower saxony basin during the late oxfordian – earliest kimmeridgian was triggered by tectonic pulse i of the late cimmerian. in the roer valley graben and the west netherlands basin, the brabant formation is locally overlain unconformably by erosional remnants of late oxfordian – portlandian deposits in continental floodplain facies (schieland group, nieuwerkerk formation). early–late kimmeridgian (stage iiib) during the kimmeridgian and portlandian, the northern dutch central graben became a major depocentre, accumulating a thick succession of increasingly marine sediments. the area of deposition once more expanded significantly. deposition resumed in continental facies in the broad fourteens basin, the dutch lower saxony basin and the vlieland basin (fig. 4). the depositional style in the broad fourteens and vlieland basins remained continental (floodplain, lacustrine, with occasional sandy fluvial intercalations), with a few marine incursions. the vlieland basin was the site of volcanic activity in the middle–late jurassic, as the zuidwal volcanic dome formed (perrot & van der poel 1987; herngreen et al. 1991). depending on the time-scale used, volcanic activity (dated radiometrically at 155–143 ma) may have taken place somewhere between the callovian and the early ryazanian. in the course of the kimmeridgian, deposition of the paralic puzzle hole formation in the central dutch central graben gradually gave way to the open marine kimmeridge clay formation in the area of blocks f08–f11. thin marine sand bodies are found locally in the uppermost parts of the puzzle hole formation, suggesting the existence of a backstepping coastal-barrier system. however, much of the character of the transgression cannot be reconstructed due to subhercynian 224 and/or laramide erosion (late cretaceous – paleocene). to the south (blocks f15–f17), the paralic puzzle hole formation grades into deposits characterised by continental alluvial plain facies (delfland formation of herngreen & wong, 1989; now called friese front formation). in the southern dutch central graben, alluvial plain deposition replaced the predominantly lacustrine deposition around the end of the oxfordian. during the early kimmeridgian (lza-4.4 of haq et al. 1988), the depositional area in the southern dutch central graben expanded. seismic and palaeogeographic information suggest that the boundaries between the areas of deposition of the puzzle hole formation, the kimmeridge clay formation and the friese front formation were determined by faults. in the broad fourteens basin, widespread deposition commenced in the early kimmeridgian (lza-4.5 of haq et al. 1988), with the accumulation of sandy alluvial plain sediments referred to the aerdenhout member of the breeveertien formation (fig. 4). this member unconformably overlies deposits of the altena group or the upper germanic triassic. the progressive transgression observed in the central graben coincides with a shift from alluvial plain (aerdenhout member) to lacustrine and lagoonal deposition (fourteens claystone member) in the broad fourteens basin (associated with the maximum flooding surface of sequence lza-4.6 of haq et al. (1988). deposition of the basal weiteveen clastic member in the dutch lower saxony basin (= niedersachsen basin) seems to have started close to the end of the early kimmeridgian. the introduction of coarser siliciclastics in the dutch central graben (main friese front member, puzzle hole formation and upper graben formation), the broad fourteens basin (breeveertien formation, aerdenhout member) and the dutch lower saxony basin (weiteveen basal clastics member) coincided with the start of the late cimmerian phase (haanstra 1963; ’t hart 1969; ‘late kimmerian i pulse’ of rgd 1991b). the resulting uplift terminated the first depositional phase of the nieuwerkerk formation in the most rapidly subsiding areas of the roer valley graben and the west netherlands basin. late kimmeridgian – earliest portlandian (stage iiic) in the late kimmeridgian, the ongoing transgression led to the first marine influence on sedimentation in the southern dutch central graben and the terschelling basin. this is witnessed by the deposition of the marineinfluenced oyster ground member (friese front formation, fig. 4). two distinct transgressive phases are identified; in the early late kimmeridgian, the transgression reached block f18, while block l02 became transgressed by the latest kimmeridgian. these events are tentatively correlated with the transgressive systems tracts of sequences lza-4.7 and lzb-1.1 of haq et al. (1988). the latter flooding event is associated with onlap onto exposed triassic and permian rocks in the terschelling basin (oyster ground member in block f15) and the vlieland basin – central netherlands basin area (continental zurich formation). at the same time, tectonic tilting and associated halokinesis caused the depocentre of the dutch central graben to shift from the central–eastern axis (from block f03, extending southwards into blocks l03–f18) to the western graben margin (block f02–l05). the seismic cross-section of the graben system shows this westwards shift (fig. 6). during the late kimmeridgian – portlandian, siliciclastic deposition in the lower saxony basin was periodically replaced by accumulation of evaporites and carbonates (evaporitic and marl members of the weiteveen formation, fig. 4). these cycles may be correlatable with the alternation of coarser clastics and fines with minor evaporites found in the broad fourteens basin (breeveertien formation, several members). the lithofacies (carbonates, marls, evaporites, coals) of the zurich formation in the remnants of the central netherlands basin (‘voorthuizen subbasin’ of haanstra 1963; gouwzee trough of rgd 1993a) are similar to those in the niedersachsen group in the dutch lower saxony basin. in the west netherlands basin and the roer valley graben, lacustrine and alluvial plain deposition was restricted to those fault blocks undergoing strongest subsidence. a progressive depositional onlap is seen in the portlandian–ryazanian section. in each consecutive depositional sequence, higher basin-fringe fault blocks became part of the depositional area. very thick successions of proximal alluvial plain deposits can be found in places. in the portlandian and ryazanian, deposition resumed in the central netherlands basin. the depositional style is transitional between the mainly lacustrine and evaporitic facies of the lower saxony basin and the paralic and restricted marine facies of the central graben – vlieland basin. deposition in the southern province (west netherlands basin and the western roer valley graben) ceased due to uplift in the kimmeridgian (around the lower sequence boundary of lzb-1.1 of haq et al. 1988). non-deposition lasted until the latest portlandian – ryazanian. from then on, continental sediments were deposited; marine conditions did not reach this area until the late hauterivian – barremian. portlandian (stage iiid) during the early portlandian, the ongoing transgression resulted in deposition of the terschelling member of the friese front formation, reflecting coastal deposition along the southern fringe of the dutch central graben (blocks l09–l12, sequences lzb-1.2, -1.3 of haq et al. 1988). to the north, the friese front formation grades into the open marine scruff greensand formation in the southern dutch central graben. thick, sand-dominated successions of portlandian–ryazanian age were deposited here (blocks f15–f18, sequences lzb-1.3 to -1.5 of haq et al. 1988). equivalent, thinner sand tongues are found in the eastern terschelling basin, the vlieland basin and the northern dutch central graben (blocks f03–f05). during this period, the vlieland basin became divided into two subbasins, separated by the zuidwal volcanic dome (herngreen et al. 1991). during the early portlandian, marine conditions only prevailed in the northern subbasin. the southern subbasin was characterised by continental (lacustrine to lagoonal) deposition. the scruff greensand formation grades into the kimmeridge clay formation in the northern dutch central graben. the depocentre of the open marine kimmeridge clay formation, which prior to the portlandian was situated in this northern area, abruptly shifted to the southernmost central graben and the northern vlieland basin. in the southern b-quadrant, basin circulation stagnated during the portlandian. this resulted in deposition of the bituminous claystones of the clay deep member. euxinic marine conditions became more widespread in the northern dutch central graben during the ryazanian (sequence lzb-1.5 and particularly lzb-1.6 of haq et al. 1988). late portlandian – ryazanian (stage iiie) during the late portlandian (sequence lzb-1.5 of haq et al. 1988), marine conditions (scruff greensand formation) briefly reached the entire vlieland basin. in the southern dutch central graben, the marine basin became shallower, inducing the northwards progradation of shallow marine, spiculitic greensands (scruff spiculite member). even the more elevated areas (for example over salt domes) of the dutch lower saxony basin and central netherlands basin, which had thus far remained exposed, became inundated by the sea. in the latter basin, the depositional area expanded to the north-west. in the broad fourteens basin, this overall regressive tendency is reflected by the transition within the breeveertien formation from lacustrine and lagoonal coastal plain fines (fourteens claystone member and driehuis mottled claystone member) to widespread sandy alluvial plain deposits (bloemendaal member). in the dutch lower saxony basin, the late portlandian shoaling trend is demonstrated by the serpulite member of the weiteveen formation. this widespread carbonate deposit marks the culmination of clastic starvation in this area. deposition of evaporites and carbonates with subordinate siliciclastics (weiteveen formation) in the lower saxony basin was replaced by open water lacustrine deposition (coevorden formation) at the beginning of the ryazanian. this event is interpreted as indicative of a flooding phase that correlated with the transition from the sandy fluvial bloemendaal member to the lagoonal neomiodon claystone member (both of the breeveertien formation) in the broad fourteens basin. this period was also marked by widespread deposition of sediments referred to the nieuwerkerk formation in the roer valley graben and west netherlands basin; thick successions of locally coarse alluvial plain sediments were deposited on rapidly subsiding fault blocks along the axes of the basins. the depositional area continued to expand gradually. the basin margins were either covered by a thin, condensed succession (e.g. locally in the province of groningen), or remained exposed and non-depositional. late ryazanian (stage iiif) during the ryazanian (sequence lzb-1.6 of haq et al. 1988), the stagnant marine basin area in the northern dutch central graben expanded southwards into block f05. deposition of the clay deep member completely superseded scruff greensand formation deposition around the mid-ryazanian. this change in depositional style coincided with the first signs of tectonic pulse ii of the late cimmerian (rgd 1991a). this also caused local truncation of the scruff greensand formation and the friese front formation in the southern central graben (parts of blocks l02–05). initially, sedimentation of the scruff greensand formation resumed briefly (stortemelk member, correlated with the lowstand of 225 sequence lzb-1.6 of haq et al. 1988). subsequently, the deep marine basin expanded markedly southwards, depositing kimmeridge clay formation sediments (schill grund member) in the previously shallow marine to continental realm. this marine incursion also reached the vlieland basin, where the stortemelk member is found intercalated with lagoonal–lacustrine deposits of the zurich formation. in the broad fourteens basin, tectonic pulse ii of the late cimmerian is expressed as two minor unconformities (base and intra-neomiodon claystone member, breeveertien formation). in the dutch central graben, large-scale differential subsidence ended with tectonic pulse ii of the late cimmerian (mid-ryazanian and earliest valanginian). from then on (sequence lzb-2.1 of haq et al. 1988), a more uniform sedimentation pattern started, associated with the post-rift thermal sag phase. this is illustrated by the contrast between the highly variable thickness of the succession underlying the stortemelk member and the more uniform development of the post-uplift succession (clay deep member, stortemelk member and schill grund member). in the broad fourteens basin, the west netherlands basin and the roer valley graben, the highly differential subsidence patterns continued at least throughout the valanginian. subsequent history the expansion of the marine sedimentation area continued during the cretaceous, until eventually the london–brabant massif became flooded in the campanian–maastrichtian. the jurassic successions in the dutch subsurface were deformed by several later tectonic phases. the santonian–campanian (subhercynian) and early paleogene (laramide) inversion phases had the most severe impact. during these compressive events, the jurassic depocentres were inverted (van wijhe 1987). practically all jurassic sediments were removed along the inversion axes of these basins. at present, the occurrences of jurassic rocks are restricted to the basinal areas (fig. 3). subsequent tertiary subsidence in the netherlands was more or less evenly distributed. subsidence rates were highest in the northern dutch offshore, however, and tertiary (particularly neogene) sediment distribution patterns are dominated by large-scale progradation in this direction. petroleum geology several jurassic formations have economic significance in the netherlands, either as reservoirs or as oil source rocks. figures 1 and 4 show the stratigraphic positions of the units mentioned in this context below, and their equivalents in neighbouring countries. the toarcian posidonia formation is the most prominent oil and wet gas source rock in the dutch subsurface (bodenhausen & ott 1981), to which the majority of the oil reserves in the country can be attributed. additional oil sourcing potential can be attributed to certain lacustrine strata of the coevorden formation (ryazanian) in the lower saxony basin. this interval is also known to act as an oil source rock in adjacent parts of germany. the highly bituminous clay deep member (ryazanian) of the kimmeridge clay formation in the northern sector of the dutch central graben has oil-sourcing potential. however, the areal distribution and stratigraphic range of this deposit is much smaller than that of equivalent bituminous intervals in the kimmeridge clay formation (and equivalents) in the uk, danish and norwegian offshore, where it constitutes one of the major oil source rocks. coal occurrences in the lower graben formation (middle–upper callovian), the middle graben formation (oxfordian), the puzzle hole formation (mainly kimmeridgian) and the friese front formation (uppermost callovian – portlandian) may locally have sourced wet gas. however, in most areas, burial was insufficient to reach the gas window (wong et al. 1989). on the other hand, gas source rocks of late carboniferous age are more or less ubiquitous throughout the netherlands. the main problem for gas generated from these older rocks is to pass through the cover of permian salts and reach the jurassic. furthermore, the maximum burial depths of the carboniferous strata, for instance in the central graben, may have caused these rocks to become overmature. at present, oil and gas reserves have been discovered in the lower graben formation (middle–upper callovian), upper graben formation (uppermost oxfordian), scruff greensand formation (portlandian) and friese front formation (uppermost callovian – portlandian) in the central graben area. some sandstone members of the breeveertien formation (kimmeridgian– ryazanian) in the broad fourteens basin were found to be oil-bearing. in the west netherlands basin, the ryazanian and younger levels of the nieuwerkerk formation are prospective. the middle werkendam member (bajocian) and carbonate members of the 226 227 f1 103 2. 0 1. 0 twt sec 3. 0 en e a lt en a g p pu zz le h o le f m . u pp er g ra be n fm pu zz le h o le f m lo w er g ra be n fm m id dl e g ra be n fm k im m er id ge c la y fm sc ru ff g re en sa nd f m . c la y d ee p m b. lo w er n o rt h se a g p z ec hs te in sa lt f m z ec hs te in sa lt f m c ha lk g p r ijn la nd g p u pp er g er m an ic t ri as g p lo w er g er m an ic t ri as g p po si do ni a sh al e fm ba se n or th s ea g ro up s ba se c ha lk g ro up ba se v al an gi ni an ba se s cr uf f g re en sa nd f or m at io n ba se k im m er id ge c la y fo rm at io n; o ys te r g ro un d m b ba se u pp er g ra be n fo rm at io n ba se m id dl e g ra be n fo rm at io n ba se l ow er g ra be n fo rm at io n ba se l ow er g er m an ic t ri as g ro up ( ba se l ow er b un ts an ds te in f or m at io n) ba se p uz zl e h ol e fo rm at io n, f ri es e fr on t fo rm at io n in tr at ri as si c le ve l i ( ne ar -b as e k eu pe r fo rm at io n) in tr at ri as si c le ve l i i (m us ch el ka lk e va po ri te m em be r) in tr at ri as si c le ve l i ii (n ea rba se m id dl e bu nt sa nd st ei n su bg ro up ) ba se /t op p os id on ia s ha le f or m at io n ba se a lte na g ro up ( ba se s le en f or m at io n) o pe n m ar in e cl ay st on e (v lie la nd c la ys to ne f or m at io n) o pe n m ar in e & la go on al c la ys to ne ( k im m er id ge c la y fo rm at io n, a lte na g ro up ) eu xi ni c re st ri ct ed m ar in e cl ay st on es ( c la y d ee p m b (k im m er id ge c la y fm ), po si do ni a sh al e fm ) o pe n m ar in e ar gi lla ce ou s sa nd st on e/ gr ee ns an d (s cr uf f g re en sa nd f or m at io n) sh al lo w m ar in e/ co as ta l s an ds to ne ( u pp er g ra be n fo rm at io n) c oa l-r ic h co as ta l p la in ( pu zz le h ol e fo rm at io n) fi ne -g ra in ed fl uv ia l p la in , o ve rb an k do m in at ed ( m id dl e g ra be n fo rm at io n) sa nd y flu vi al p la in , c ha nn el d om in at ed ( lo w er g ra be n fo rm at io n) r ef le ct o rs d ep o si ti o n al s et ti n g/ se is m ic fa ci es f0 802 f0 901 w sw en e fi g. 6 . s ei sm ic c ro ss -s ec tio n , r u n n in g w sw –e n e a cr o ss t h e d u tc h s ec to r o f th e c en tr al g ra b en ( fi g. 2 ). t h e ju ra ss ic – l o w er c re ta ce o u s se ct io n h as b ee n c o lo u rco d ed a cc o rd in g to t h e m ai n d ep o si tio n al f ac ie s. brabant formation (callovian–oxfordian) have locally been shown to be oil-bearing in the roer valley graben and the west netherlands basin. exploration of the jurassic in the netherlands has reached a mature stage. the acquisition of high-quality seismic and well data and the good biostratigraphic control obtained in recent years have resulted in a clear regional geological picture. these data also allow, for the first time, the application of sequence stratigraphy. integration of sequence stratigraphic concepts into the revision of dutch lithostratigraphy (van adrichem boogaert & kouwe 1994–1997) has resulted, in particular, in a far better understanding of the upper jurassic stratigraphy of the netherlands. acknowledgements the authors wish to thank the geological survey of the netherlands (rijks geologische dienst – rgd/nitg) for permission to publish the presented data. the authors are indebted to r. rijkers, a. hollen and f. rispens (all rgd) for providing information on the regional distribution and petroleum exploration successes in the dutch jurassic. h.a. van adrichem boogaert and m.c. geluk (rgd) gave valuable comments to the manuscript and its precursors. similarly, the manuscript has benefited significantly from earlier discussions with the various members of the working group on the upper jurassic – lower cretaceous stratigraphy for the new stratigraphic nomenclature of the netherlands. references bodenhausen, j.w.a. & ott, w.f. 1981: habitat of the rijswijk oil province, onshore, the netherlands. in: illing, l.v. & hobson, g.d. 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(compiler) 1974: a proposed standard lithostratigraphic nomenclature for the southern north sea and an outline structural nomenclature for the whole of the (uk) north sea. institute of geological sciences report 74/8, 14 pp. london: her majesty’s stationery office. ’t hart, b.b. 1969: die oberjuraund unterkreide-sedimentation in den nördlichen und östlichen niederlanden. erdöl und kohle, erdgas, petrochemie 22, 253–261. underhill, j.r. & partington, m.a. 1993: jurassic thermal doming and deflation in the north sea: implications of the sequence stratigraphic evidence. in: parker, j.r. (ed.): petroleum geology of northwest europe: proceedings of the 4th conference, 337–345. london: geological society. van adrichem boogaert, h.a. & kouwe, w.f.p. (compilers) 1994–1997: stratigraphic nomenclature of the netherlands, revision and update by rgd and nogepa. mededelingen rijks geologische dienst 50, sections a–j (sections paginated independently). van wagoner, j.c., posamentier, h.w., mitchum, r.m., vail, p.r., sarg, j.f., loutit, t.s. & hardenbol, j. 1988: an overview of the fundamentals of sequence stratigraphy and key definitions. in: wilgus, c.k. et al. (eds): sea-level changes – an integrated approach. society of economic paleontologists and mineralogists special publication 42, 39–45. van wijhe, d.h. 1987: structural evolution of inverted basins in the dutch offshore. in: ziegler, p.a. (ed.): compressional intraplate deformations in the alpine foreland. tectonophysics 137, 171–219. wong, th.e., van doorn, th.h.m. & schroot, b.m. 1989: ‘late jurassic’ petroleum geology of the dutch north sea central graben. geologische rundschau 78, 319–336. ziegler, p.a. 1990: geological atlas of western and central europe, 2nd edition, 239 pp. amsterdam: elsevier for shell internationale petroleum maatschappij. 229 manuscript received 11 september 1995; revision accepted 17 january 1997. gsb kimmeridge clay fm cft cfr m. graben fm l. graben fm clay deep mb scruff greensand fm friese front fm puzzle hole fm cfo schill grund mbgss gsp gsakim. clay fm zurich fm fourteens claystone mb santpoort mb bloemendaal mb aerdenhout mb coevorden fm a b c e d f n ie de rs ac hs en g p sc hi el an d sc ru ff g ro up kn kn kn neomiodon mb r ya z. valang. k im . k im m er id gi an t ith on ia n po rt la nd ia n be rr ia s. m al m t et hy s bo re al d og ge r li as α γ−ζ ζ ε7 ε6 ε5 δ2 δ1 α δ2 β1 β2-4 ζ ε2 δ1 γ3 β3 γ2 β2 γ1 β1 α3 α2 α1 ε1 γ ε1-4 oxford. callov. g lo ba l su bs ta ge s se qu en ce s af te r h aq e t a l. (1 98 8) ba th on ia n ba jo ci an a al en ia n t oa rc ia n si ne m ur ia n h et ta ng ia n pl ie ns ba ch ia n norian rhaetian 157.1 154.7 152.1 145.6 140.7 147.8 142.8 time (ma) after harland et al. (1990) stratigraphy 203.5 208.0 209.5 194.5 187.0 173.5 178.0 166.1 161.3 weiteveen fm breeveertien fm nieuwerkerk fm nieuwerkerk fm roer valley graben s n west netherlands basin central netherlands basin achterhoek lower saxony basin dutch central grabenbroad fourteens basin aalburg fm sleen fm posidonia shale fm werkendam fm ? ? ? lower werkendam claystone mb upper werkendam claystone mb brabant fmatbr1 atbrl atbr2 atbrm atbr3 atbro klomps mb atbru middle werkendam mb β sc hi el an d g ro up a lte na g ro up rn u a b3 u a b4 lz a -1 lz a -2 lz a -3 lz a -4 lz b1 uab-2 uab-1 uaa-4 lzb-2.1 sc hi el an d g ro up sc hi el an d g p cms u. graben fm driehuis mb depositional facies marl open marine clay (cretaceous) open marine clay (jurassic) bituminous marine clay/shale open marine argillaceous greensand shallow marine greensand coastal sand coal-rich coastal plain heterolithics sandy coastal/fluvial plain heterolithics fine-grained coastal plain and lacustrine heterolithics lagoonal claystones/siltstones lacustrine/highly restricted marine carbonates highly restricted marine salts highly restricted marine anhydrite/carbonates shallow marine sandy limestone abbreviated stratigraphic terms brabant formation atbro: oisterwijk limestone mb atbru: upper brabant marl mb atbr3: upper brabant limestone mb atbrm: middle brabant marl mb atbr2: middle brabant limestone mb atbrl: lower brabant marl mb atbr1: lower brabant limestone mb middle graben formation cms: middle graben sandstone mb friese front formation cfo: oyster ground claystone mb cft: terschelling sandstone mb cfr: rifgronden claystone mb weiteveen formation f: serpulite mb e: upper marl mb d: upper evaporite mb c: lower marl mb b: lower evaporite mb a: basal clastic mb scruff greensand formation gss: stortemelk mb gsp: scruff spiculite mb gsa: scruff argillaceous mb gsb: scruff basal sandstone mb kn: rijnland gp rn: upper germanic trias gp 1.6 1.5 1.4 1.3 1.2 1.1 4.7 4.6 4.5 4.4 4.2 4.1 3.2 3.1 2.4 2.3 2.2 2.1 1.1 4.6 4.5 4.4 4.3 4.2 4.1 3.4 3.3 3.2 3.1 4.3 fig. 4. litho-chronostratigraphic scheme of the rhaetian–ryazanian succession of the five main jurassic basin systems in the netherlands: (1) roer valley graben – west netherlands basin, (2) off-holland low – broad fourteens basin, (3) central netherlands basin, (4) ems low – lower saxony basin and (5) vlieland basin, terschelling basin and dutch central graben. abbreviated stage names are given in full in fig. 1; berrias., berriasian. compiled and modified from van adrichem boogaert & kouwe (1994–1997, sections f, g). lza-1 uab-4 uab-3 uab-2 u a b1 u a a -41. 1 4. 6 4. 5 4. 4 4. 3 4. 2 3. 4 3. 3 3. 2 3. 1 2. 1 4. 1 4. 4 lza-3 lza-2 3. 2 3. 1 2. 3 2. 2 2. 1 2. 4 4. 1 4. 2 4. 3 4. 5 el m oxfordian callovian el m sequences (haq et al. 1988) lza-4lzb-1lzb-2 ryazanian valanginian 2. 3 kimm. kimmeridgian s.l. el tithonian portlandian berriasian 4. 6 4. 7 1. 1 1. 2 1. 3 1. 4 1. 5 1. 6 ‘e’e ‘l’l 2. 1 2. 2 el st ra tig ra ph y 13 5 14 0 14 5 15 0 15 5 16 0 16 5 17 0 17 5 18 0 18 5 19 0 19 5 20 0 20 5 21 0 t im e (m a) doggermalm lias ζ ε δ 2 δ 1 α δ 2 β 2– 4 ζ ε 2 δ 1 γ 3 β 3 γ 2 β 2 γ 1 β 1 α 3 α 2 α 1 ε 1γε 1– 4 bathonian bajocian aalenian toarcian sinemurian hettangianpliensbachian n or ia n triassicearly jurassicmiddle jurassiclate jurassiccretaceous r ha et ia n β 1 e l e l eml d ic ho to m ite s un na m ed n w e ur op ea n st an da rd a m m on ite z on at io n po ly pt yc hi te s pa ra to lli a al bi du m st en om ph al us ic en ii ru nc to ni la m pl ug hi pr ep lic om ph al us pr im iti vu s ‘o pp re ss us ’/a ng ui fo rm is ke rb er us ok us en si s gl au co lit hu s al ba ni fit to ni ro tu nd a pa lla si oi de s pe ct in at us hu dl es to ni w he at le ye ns is sc itu lu s el eg an s au tis si od or en si s eu do xu s m ut ab ili s cy m od oc e ba yl ei se rr at um /g lo se ns e te nu is er ra tu m de ns ip lic at um co rd at um /m ar ia e la m be rt i at hl et a co ro na tu m ja so n ca llo vi en se m ac ro ce ph al us di sc us or bi s ho ds on i m or ri si su bc on tr ac tu s pr og ra ci lis te nu ip lic at us zi gz ag pa rk in so ni ga ra nt ia na ni or te ns e hu m ph ri es ia nu m pr op in qu an s la ev iu sc ul a di sc ite s co nc av um br ad fo rd en si s m ur ch is on ae op al in um aa le ns is ps eu do ra di os a di sp an su m th ou ar se ns e va ri ab ili s bi fr on s se rp en tin us te nu ic os ta tu m sp in at um da vo ei ib ex ja m es on i ra ri co st at um se m ic os ta tu m bu ck la nd i an gu la ta lia si cu s pl an or bi s m ar sh i su es si m ac er ox yn ot um ob tu su m tu rn er i m ar ga ri ta tu s ro se nk ra nt zi /r eg ul ar e ko ch i se le ct ed s po ro m or ph d at um s va ri ou s tr ilo bo sp or ite s sp p. cl av ife ra tr ip le x ae qu itr ira di te s; pl ic at el la cl as so po lli s cl as so po lli s cl as so po lli s ec hi na tu s/ ha m m en ii pl ex us c ci ca tr ic os isp or ite s pr ec ic at ric os isp or ite s pr ec ic at ric os isp or ite s sp p. ca lli as po rit es s eg m en ta tu s; ly co po di ac id ite s ru gu la tu s; n eo ra ist ric ki a tr un ca ta ; u va es po rit es a rg en te ae fo rm is; c ly co po di um sp or ite s se m im ur is ca lli as po rit es s eg m en ta tu s ch as m at os po rit es m aj or ci rc ul in a m ey er ia na tr ili te s sp p. is ch yo sp or ite s va rie ga tu s; st ap lin isp or ite s (in cl . c or on at isp or ite s) c ro ga lsk ai sp or ite s ci ca tr ic os us ca lli as po rit es s pp .; n eo ra ist ric ki a tr un ca ta le pt ol ep id ite s sp p. ; s es tr os po rit es p se ud oa lve ol at us c h el io sp or ite s al tm ar ke ns is co nt ig ni sp or ite s pr ob le m at ic us ce re br op ol le ni te s sp p. o va lip ol lis p se ud oa la tu s ci ng ul iz on at es r ha et ic us ; c or nu tis po rit es s ee be rg en sis ; d en so sp or ite s fis su s; lu na tis po rit es r ha et ic us ; rh ae tip ol lis g er m an ic us ; s em ire tis po ris s pp .; tr ia nc or ae sp or ite s re tic ul at us ; ts ug ae po lle ni te s ps eu do m as su la e; z eb ra sp or ite s la ev ig at us ri cc iis po rit es tu be rc ul at us ; z eb ra sp or ite s in te rs cr ip tu s n eo ra ist ric ki a gr ist ho rp en sis ; u va es po rit es ar ge nt ea ef or m is; c d en so isp or ite s; se st ro sp or ite s ps eu do al ve ol at us ra ist ric ki sp or ite s br ev itr un ca tu s c co nc av iss im isp or ite s; c im pa rd ec isp or a; c tr ilo bo sp or ite s; kr ae us el isp or ite s sp . pa rv isa cc ite s ra di at us kr ae us el isp or ite s tu bb er ge ns is m . e xt en siv a se le ct ed d in of la ge lla te d at um s t. a pa te la e. p ha ro t. d av ey i b. r ad ic ul at um g . v irg ul a c. p an ne um ; g . d im or ph um g . m ut ab ili s; s. ju ra ss ic a c m ud er on gi a sp . a s. in rit ib ilu m a. d ic ty ot a; p . p an no su m p. p an no su m n . p el lu ci da ; s . c ry st al lin um l. s ca rb ur gh en sis li th od in ia ju ra ss ic a; c pa re od in ia p ro lo ng at a n an no ce ra to ps is sp ic ul at a; n . g ra ci lis l. s ca rb ur gh en sis c r. a em ul a; w an ae a c. c on tin uu m ca rp at ho di ni um p re da e be ju ia p ol yg on al is n an no ce ra to ps is di ct ya m bo ni s; n . t ric er as ; w al lo di ni um c yli nd ric um sc rin io ca ss is w eb er i lu eh nd ea s pi no sa li as id iu m v ar ia bi le d ap co di ni um p ris cu m rh ae to go ny au la x sp p. pa rv oc ys ta s pp .; re ut lin gi a sp p. ; su sa di ni um s pp .; c ph al lo cy st a eu m ek es m ei ou ro go ny au la x va le ns ii rh yn ch od in io ps is re ga lis ; v al va eo di ni um s pi no su m ct en id od in iu m c om ba zi i/o rn at um p le xu s st ep ha ne lyt ro n; s . f as c. /p en ic . c p. p an no su m e. lu rid um r g . j ur as sic a; r o . p at ul um o . p at ul um c cr ib ro pe rid in iu m s p. a /b ; e . p ol yp la co ph or um g on ya ul ac ys ta s p. a ; r . t hu la d . s pi no su m ; e . t or yn um ; k . p or os isp in um c p. ‘e op el lif er um ’ s en su r r i ( 19 87 ) am m ov er te lla c el le ns is (f ); pr ot oc yt he re h an no ve ra na ( o) ; pr ot oc yt he re p se ud op ro pr ia ( o) m an de lst am ia s ex ti (o ); cy th er op te rin a tr ie be li (o ); g al lia ec yt he rid ea te re s (o ); pa ra no ta cy th er e sp ee to ne ns is (o ); sc hu le rid ea ju dd i ( o) cy pr id ea c ar in at a (o ); c. g ra nu lo sa ( o) g al lia ec yt he rid ea p os ts in ua ta ( o) ; k lie an a di ct yo ta ( o) fa ba ne lla a ns at a (o ); m an te lli an a pu rb ec ke ns is (o ) pa ra no ta cy th er e el on ga ta ta ( o) ; g al lia ec yt he rid ea c om pr es sa ( o) g al lia ec yt he rid ea p ol ita ( o) ; p ar al es le ya p er fo ra ta ( o) g al lia ec yt he rid ea s pi no sa ( o) ; m an de lst am ia tu m id a (o ); sa ra ce na ria o xf or di an a (f ); aa le ni el la in or na ta ( o) eo cy th er id ea e us ar ca ( o) ; g al lia ec yt he rid ea v ol ga en sis ( o) ; m ac ro de nt in a ru gu la ta ( o) ; m ac ro de nt in a tr an sie ns ( o) ; ex op ht ha lm oc yt he re g ig an te a (o ); kl ie an a ca lyp tr oi de s (o ) se le ct ed fo ra m in ife r/ os tr ac od d at um s m an de lst am ia m ac ul at a (o ); ep ist om in a nu da ( f) m an de lst am ia r ec til in ea ( o) ve rn on ie lla s eq ua na ( o) lo ph oc yt he re m ul tic os ta ta ( o) n op hr ec yt he re c ru ci at a ox fo rd ia na ( o) ps eu do pe ris so cy th er id ea p ar ah ie ro gl yp hi ca ( o) ; eu cy th er ur a co st ae irr eg ul ar is (o ) g al lia ec yt he rid ea p un ct at a (o ); m ac ro de nt in a ci ca tr ic os a (o ); er ip le ur a el ea no ra e (o ); g al lia ec yt he rid ea d iss im ili s (o ) sc hu le rid ea tr ie be li (o ); g al lia ec yt he rid ea w ol bu rg i ( o) eu cy th er ur a ho rr id a (o ) lo ph oc yt he re fl ex ic os ta ( o) ; n op hr ec yt he re c ru ci at a cr uc ia ta ( o) ; ep ist om in a m os qu en sis ( f) lo ph oc yt he re in te rr up ta in te rr up ta ( o) lo ph oc yt he re s ca br a sc ab ra ( o) ; l . b ip ar tit a (o ); ps eu do hu ts on ia tu be ro sa ( o) pr og on oc yt he re s til la ( o) ; g lyp to cy th er e gu em be lia na ( o) pr og on oc yt he re p ol on ic a (o ); lo ph oc yt he re p le na ( o) pl eu ro cy th er e co nn ex a (o ); le nt ic ul in a qu en st ed ti (f ) g lyp to cy th er e au ric ul a (o ) am m op al m ul a in fr aj ur en sis ( f) ; f uh rb er gi el la g ig an te a (o ); pl eu ro cy th er e im pa r (o ); g lyp to cy th er e tu be ro de nt in a (o ) pl eu ro cy th er e re gu la ris ( o) fu er be rg ie lla p rim iti va ( o) g lyp to cy th er e sc itu la ( o) ; f uh rb er gi el la h or rid a ho rr id a (o ); lj ub im ov el la p iri fo rm is (o ) g lyp to cy th er e po lit a (o ) pr ae sc hu le rid ea d ec or at a (o ); ca m pt oc yt he re p us ill a (o ) ca m pt oc yt he re m ed ia ( o) ca m pt oc yt he re fo ve ol at a (o ) o to cy th er e ca llo sa ( o) ; a ph el oc yt he re k uh ni ( o) ; ca m pt oc yt he re p ra ec ox ( o) ; l en tic ul in a fo ve ol at a (f ) ki nk el in el la p er sic a (o ) li ng ul in a te ne ra g ro up ( f) ; m ar gi nu lin a pr im a gr ou p (f ) fr on di cu la ria te rq ue m i ( f) pl eu rif er a ha rp a (o ); g am m ac yt he re u bi qu ita ( o) ; o gm oc on ch el la b isp in os a (o ); g ra m an ni cy th er e ba ch i ( o) ; o gm oc on ch a am al th ei ( o) ; b ai rd ia c lio ( o) ; p ol yc op e ci nc in na ta ( o) kl in gl er el la v ar ia bi lis ( o) lo ph od en tin a pu m ic os a (o ); ek ty ph oc yt he re b et zi ( o) ; ek ty ph oc yt he re tr ie be li (o ); lo ph od en tin a la cu no sa ( o) o gm oc on ch a ha ge no w i ( o) ; n an ac yt he re e le ga ns ( o) o gm oc on ch el la e lli ps oi de a (o ); cy th er el lo id ea c irc um sc rip ta ( o) ; cy th er el lo id ea p ul ch el la ( o) ; k lin gl er el la tr an slu ce ns ( o) rh om bo cy th er e pe na rt he ns is (o ) cy th er el lo id ea b ui se ns is (o ); li ng ul in a te ne ra c ol le no ti (f ); lu tk ev ic hi ne lla s p. ( o) o gm oc on ch a co nt ra ct ul a (o ); tr ac hy cy th er e tu bu lo sa ( o) ; sa ra ce na ria s ub la ev is (f ); bo liv in a lia sic a lia sic a (f ) w ic he re lla s em io ra s em io ra ( o) ; g ra m an ne lla a po st ol es cu i ( o) m ic ro pn eu m at oc yt he re s ub co nc en tr ic a (o ); o lig oc yt he re is fu llo ni ca ( o) pl eu ro cy th er e ric ht er i ( o) to p ac m e ba se a cm e co ns is te nt ra re fo ra m in ife r os tr ac od sp or om or ph a pp ea ra nc es a nd di sa pp ea ra nc es a re a pp ro xi m at e to p oc cu rr en ce ba se o cc ur re nc e (f ) (o ) c r fi g. 5 . a s el ec tio n o f d ia gn o st ic b io m ar ke r h o ri zo n s fo r th e r h ae tia n –v al an gi n ia n , co rr el at ed w ith t h e n o rt h w es t e u ro p ea n s ta n d ar d a m m o n ite z o n at io n . t h e h aq e t a l. (1 98 8) c yc le c h ar t h as b ee n r ec al ib ra te d t o f it th e tim esc al e o f h ar la n d e t a l. (1 98 9) . k im m er id gi an s .l. eq u al s k im m er id gi an s en su a n gl ic o. t h e te rm s li as , d o gg er a n d m al m a n d t h ei r cl as si c re gi o n al s u b d iv is io n s as d ep ic te d i n t h e va ri o u s fi gu re s ar e u se d i n o rd er t o s h o w t h e st ra tig ra p h ic p o si tio n o f th e se d im en ts f o u n d i n t h e n et h er la n d s in a e u ro p ea n c o n te xt o f w el les ta b lis h ed z o n atio n s. c o m p ile d a n d m o d if ie d f ro m v an a d ri ch em b o o ga er t & k o u w e (1 99 4– 19 97 , se ct io n s f, g ). geological survey of denmark and greenland bulletin 26, 2012, 57-60 57 from 3d mapping to 3d modelling: a case study from the skaergaard intrusion, southern east greenland kristian svennevig and pierpaolo guarnieri the powerful 3d mapping tool at the photogrammetry laboratory of the geological survey of denmark and greenland (geus) is ideal for collecting high-quality 3d geological data in remote and inaccessible areas with a high degree of exposure such as greenland (vosgerau et al. 2010). so far this 3d mapping tool has been used to visualise and extract very precise geological data from aerial and oblique photographs. in the study reported on here, the 3d mapping tool was used to generate data for 3d geological modelling. the skaergaard intrusion (fig. 1) is a well-known eocene layered gabbro. the study of the intrusion has had great importance for the understanding of magmatic petrology, magma differentiation and fractional crystallisation since the early studies by wager & deer (1939). it was chosen for 3d modelling because it is well studied from a petrological point of view and because the shape of the magma chamber was previously modelled in a network of 2d cross sections (nielsen 2004). in this paper, it is modelled for the first time in 3d using a detailed 1:20 000 scale geological map (mcbirney 1989), 1:27 000 scale aerial photographs from 1973, data from drill holes and geophysical data. 3d mapping the 1:27  000 scale aerial photographs where scanned, triangulated and georeferenced with socketset software. the geological map was digitised as well as georeferenced using 1:150 000 scale aerial photographs and finally draped on a high-resolution (10 × 10 m grid) digital elevation model derived from the 1:27 000 scale aerial photographs. geological features were drawn as polylines from both the aerial photographs and the geological map. visible parts of the outer boundary were collected using the 3d stereoplotter (fig. 2a) with very high accuracy (vosgerau et al. 2010). oblique aerial photos were not used for 3d-mapping as the area is not sufficiently covered and the setting up of the few available photo lines is beyond the scope of this paper. other boundaries defined from mineral parageneses and compositions (wager & brown 1968) were digitised from the 1:20 000 map in arcgis. the boundaries from the aerial photographs are much more accurate than those of the map (fig. 3) because they represent directly observable geological features. furthermore polylines from the aerial photographs have real 3d coordinates measured directly on the 3d stereo plotter, whereas polylines from the arcgis map have x-y coordinates from the map and z coordinates imported from the digital elevation model. this leads to additional uncertainty as elevations in the arcgis map and the digital elevation model are not generated from the same source and sometimes do not overlap in the entire area of interest. structural readings (dip direction and dip angle) for the geological boundaries were generated from polylines using the in-house-developed software tool geus planes (fig. 2b). this tool calculates strike, dip and standard deviations on strike and dip for each node (vertex) of a polyline, generating large amounts of strike-dip measurements from the 3d geometry of the geological boundaries. 3d modelling and discussion the 3d modelling of the skaergaard intrusion was carried out using the software tool leapfrog-3d. this software can generate surfaces and volumes from structural readings, fig. 1. the skaergaard intrusion area, southern east greenland. skaergaard intrusion mainly palaeogene volcanic rocks palaeogene gabbro, including the skaergaard intrusion mesozoic precambrian basement greenland 5 km miki fjord watkins fjord glaciers fig. 3 31°25´w 68°15´n © 2012 geus. geological survey of denmark and greenland bulletin 26, 57–60. open access: www.geus.dk/publications/bull 5858 a basement skaergaard intrusion 31°43´w 1 km 68°11´n b fig. 3. example of a polyline (blue) representing the boundary between the skaergaard intrusion (dark grey in fig. 3a, light yellow to red in fig. 3b, representing different lithological units) and the basement (light grey in fig. 3a, light skin tone in fig. 3b). the same polyline is drawn on the 1:27 000 aerial photographs (a) and shown on the geological map of mcbirney (1989) (b) for comparison. the difference between the locations of the two lines is up to 100 m. fig. 2. simplified workflow diagram describing the process of generating 3d data from photographs to 3d modelling. a: polylines representing geological boundaries or structures are extracted from a 3d stereo plotter. b: the nodes of the polylines are converted into structural readings (here shown as small blue discs) using the geus planes tool. measurements with a poorly defined plane (high standard deviation on either the dip, dip-azimuth or plane) are discarded. c: the structural readings are imported into a 3d modelling software (leapfrog 3d) where surfaces and volumes can be generated and compiled to a 3d model of the geological object. a b c 59 points and polylines (fig. 2c). in standard surface modelling, only a couple of structural readings are used for generating surfaces. the great number of structural readings produced by geus planes present a problem to the modelling software making it necessary to manually select which strike-dip measurements to use for the construction of surfaces. this complication will be solved in a forthcoming upgrade of the geus planes tool by changing the output and by using modelling software that can import the standard deviation as a parameter and use it to estimate the quality of the structural readings in the creation of surfaces. in areas with no pre-quaternary exposures, such as glaciated areas or water-covered areas, support points were introduced to guide the surfaces. drill-hole data were also imported along with a geophysical gravity model of the floor of the intrusion (t.m. rasmussen, personal communication 2011) to further improve the model. due to erosion and lack of exposures the locations of the roof and northern boundary fig. 4. 3d model of the skaergaard intrusion in e–w (a) and n–s (b) sections. the view is towards 320° at a plunge of 15°. no vertical exaggeration. notice the wedge-like shape of the hidden zone (black arrow in a). 5 km a b elevation east east elevation ? upper border series basistoppen sill upper zone triple group lower zone middle zone hidden zone l ay er ed s er ie s marginal border series 6060 of the intrusion are poorly constrained, and these were modelled with considerable uncertainty. two structural readings (strike/dip, 180/20 s and 180/10 s) were constructed to guide the roof surface of the magma chamber. these orientations were estimated by extrapolating the roof boundary in the southern part of the intrusion (wager & deer 1939; nielsen 2004) with the regional dip (10–20°s) of the tertiary strata hosting the intrusion (nielsen & brooks 1981). the geophysical model for the floor of the intrusion has the same dip as the regional strata of 10–20°s but also shows a dip of around 15° towards the east that is unexpected from the surface and drill-hole data used for generating the model. this is reflected in the wedge-like shaped of the hidden zone in the 3d model (fig. 4a). furthermore, the modelled floor of the intrusion near the northern boundary is located much more deeply than expected. based on a single drill hole in the northern part of the intrusion that reaches the contact zone of the floor (holness et al. 2007), unexposed gabbro of the intrusion is estimated to reach a depth of no more than 450 m (t.f.d. nielsen, personal communication 2012), whereas the geophysical model predicts 1450 m. several explanations can be suggested for this discrepancy: (1) the density of the gabbro at the northern margin may be underestimated, (2) uncertainty of the floor depth from drill-hole data or (3) rock types below the floor of the intrusion may have unexpected high densities. irrespective of the various uncertainties, the preliminary estimate of the total volume of the intrusion is between 279 and 305 km3 (depending on different estimates of the northern boundary). this is in good agreement with a previous estimate of 280 ± 23 km3 (nielsen 2004). the triple group (fig. 4), defined by three distinctive leucocratic layers in the uppermost part of the middle zone, is of particular interest as it is rich in palladium and gold (bird et al. 1991; nielsen et al. 2005). the degree of detail in the model presented here is too low for a thorough investigation of these levels, but our data reflect the concave nature of the layers (fig. 4) also described by nielsen (2004). future expansion of the model with data from oblique photographs can greatly improve our understanding of the triple group, because it is clearly seen on photographs (fig. 11c in brooks 2011), being well exposed on the steep mountain slopes. concluding remarks the 3d mapping tool at the geus photogrammetry laboratory is ideal for generating high quality data for 3d modelling. it is, however, necessary to fine-tune the geus planes tool and to conduct the modelling in a more advanced software such as gocad or move in order to take full advantage of the high quality and quantity of the data. with these software tools, it is also possible to do advanced modelling and, for example, carry out balanced unfolding and backstripping of geological structures and units. acknowledgements max nykjær strunck is thanked for setting up the aerial photographs and for helping with the digital elevation model. thorkild maack rasmussen is thanked for producing the geophysical model and troels f.d. nielsen is thanked for valuable discussion of the geology of the skaergaard intrusion. references bird, d.k., brooks, c.k., gannicott, r.a. & turner, p.a. 1991: a goldbearing horizon in the skaergaard intrusion, east greenland. economic geology 86, 1083–1092. brooks, c.k. 2011: the east greenland rifted volcanic margin. geological survey of denmark and greenland bulletin 24, 96 pp. holness, m.b., nielsen, t.f.d. & tegner, c. 2007: textural maturity of cumulates: a record of chamber filling, liquidus assemblage, cooling rate and large-scale convection in mafic layered intrusions. journal of petrology 48, 141–157. mcbirney, a.r. 1989: geological map of the skaergaard intrusion, east greenland. eugene, usa: university of oregon (map sheet). nielsen, t.f.d. 2004: the shape and volume of the skaergaard intrusion, greenland: implications for mass balance and bulk composition. journal of petrology 45, 507–530. nielsen, t.f.d. & brooks, c.k. 1981: the east greenland rifted continental margin: an examination of the coastal flexure. journal of the geological society (london) 138, 559–568. nielsen, t.f.d., andersen, j.c.ø. & brooks, c.k. 2005: the platinova reef of the skaergaard intrusion. in: mungal, j.e. (ed.): exploration for platinum group element deposits. mineralogical association of canada short course series 35, 431–455. vosgerau h., guarnieri p., weibel r., larsen m., dennehy, c., sørensen, e.v. & knudsen, c. 2010: study of a palaeogene intrabasaltic sedimentary unit in southern east greenland: from 3-d photogeology to micropetrography. geological survey of denmark and greenland bulletin 20, 75–78. wager, l.r. & deer, w.a. 1939 (re-issued 1962): geological investigations in east greenland. part iii. the petrology of the skaergaard intrusion, kangerdlugssuaq, east greenland. meddelelser om grønland 105(4), 352 pp. wager, l. r. & brown, g.m. 1968: layered igneous rocks, 588 pp. edinburgh: oliver and boyd. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ksv@geus.dk geological survey of denmark and greenland bulletin 41, 2018, 95-98 95 geology does not respect national borders. hence, in order to get geological overviews of europe, input from geological surveys in more than 35 countries is required. european policy makers have several times been forced to rely on the us geological survey to provide e.g. resource estimates from the european continent, but for obvious reasons there is a wish to base european decision making on european knowledge. consequently, the european commission and the european parliament have formulated a request for the establishment of a ‘geological service for europe’. in its strategy towards 2020, eurogeosurveys (egs) addresses the creation of such a service through three pillars. egs is an umbrella organisation through which national geological survey organisations of 36 european countries cooperate, referred to below as national surveys. the three pillars are designed to integrate input from all national surveys into a system that can swiftly act on urgent needs for knowledge-based decision support. the three pillars relate to joint research, data integration and sharing of facilities (fig. 1). whilst the third pillar has only recently been dealt with, the two first have already advanced through a number of recent initiatives. having been a key player in numerous eu projects for many years, the geological survey of denmark and greenland (geus) has attained a central role in the implementation of these two pillars of the strategy, both as coordinator of the european geological data infrastructure (egdi, www.europe-geology.eu) and as one of the biggest players in the so-called geoera programme. geus participates in ten projects and is a partner in the secretariat and the coordinator of the geoera information platform. the present paper outlines the main steps towards the current situation and provides a background for geus’ role in this. towards a eu geological knowledge base the european commission has contributed financially through several framework programmes to increase knowledge sharing, capacity building as well as cross-border and pan-european research within different geoscience domains. in most cases, data play a central role, and egs members have many years of experience in working together with the purpose of making geological data fair (findable, accessible, interoperable and reusable). in many cases, european projects have historically only had the participation of a limited number of national surveys, but in 2008 a flagship project funded by the eu was launched with the participation of 20 national surveys as a logical follow-up of the global onegeology initiative (www.onegeology.org). onegeology aimed at assembling a geological map of the world on a scale of 1:1 000 000 by piecing together national contributions in whatever format they had, using distributed web services. the onegeologyeurope project went a step further. geus and the 19 other european national surveys worked together for two years to produce a distributed, web-based surface geological map of europe on a scale of c. 1:1 000 000 that was harmonised with a common data structure and agreed geological classifications. at that time, the project represented a leading edge activity, as it demonstrated the power of national organisations working together around a common data structure and nomenclature for geological units (geochronology and lithology). importantly, the project built on principles and rules from the inspire directive from 2007 and was thus not only a step on the way for the national surtowards a common geological data infrastructure for europe jørgen tulstrup and mikael pedersen the geological surveys of europe jo in t r es ea rc h (g eo er a) da ta in te gr ati on a nd ha rm on is ati on (e g di ) kn ow le dg e, c ap ac ity an d in fr as tr uc tu re s ha rin g european geological service fig. 1. the three pillars of the egs strategy for the establishment of a european geological service. © 2018 geus. geological survey of denmark and greenland bulletin 41, 95–98. open access: www.geus.dk/bulletin http://www.europe-geology.eu http://www.onegeology.org http://www.geus.dk/bulletin 9696 veys to fulfil their obligations towards the directive, but also served to demonstrate the feasibility of making geological data interoperable for europe and beyond. the results of the project were fed into the legal process to define the eu-wide inspire geological data specification, which since 2014 constitutes the implementing rules. upon the success of onegeologyeurope, a natural extension came with the increased attention on securing the supply of critical raw materials for european industry. this was partly based on the rare-earth element trade dispute that began in 2010 when china imposed strict export quotas for rare-earth elements. such elements are used in a number of high-technology industries, and since china accounts for 97% of the world production, the situation was considered critical (kalvig & machacek 2018, this volume). consequently, the european commission urgently needed an overview of raw materials resources in europe. this led to a number of eu projects. especially one of them had strategic importance for the egs, namely the minerals4eu project. geus cooperated with 27 other national surveys and other eu organisations to build the foundation for a european raw materials knowledge base, by extending the onegeologyeurope philosophy and by complying with and contributing to inspire. subsequently, other eu projects like eurare, prosum, mica and the recently launched orama project have extended and/or improved this common knowledge base. at the same time geus also participates in the knowledge and innovation community (kic) for raw materials, through which a number of complementary projects are run in cooperation with universities and industrial partners. simultaneously with these raw materials projects, a number of other european data harmonisation projects were carried out within other egs research areas such as groundwater, energy, geohazards and soil. however, it became increasingly clear along the way that there was a need for coordination in order to increase the efficiency, reusability and sustainability – not only to meet european expectation, but also for the sake of geoscientists. this was addressed in the egs strategy that was published in 2014 and laid the foundation for the european geological data infrastructure. the european geological data infrastructure in 2012, egs’ members were granted a two-year eu project called egdi-scope, aiming to assess the possibilities of setting up a long-term sustainable european geological data infrastructure in line with the second pillar of the egs strategy. the initiative addressed the fact that almost all previous common european geoscience projects had succeeded in producing good european datasets and commonly also web-based dissemination platforms, but that such systems would typically disappear after a certain period because of lack of financial support for basic operation and maintenance. the egdi-scope project revealed that data from more than 80 past european projects, worth 400–700 million euro, could potentially be ‘saved’ and made available through a common data infrastructure which should also be the natural dissemination platform for future projects. geus was part of the egdi-scope core team and was in charge of stakeholder consultation, thereby analysing both end-user needs and interfaces to other large european earth science projects and initiatives. an important goal of egdi-scope was to pave the way for a larger eu project, whereby egdi could be implemented. fig. 2. the interactive map viewer of the european geological data infrastructure (egdi) portal through which hundreds of different data themes can be viewed in combination. the map shown here illustrates the concentration of lead in grazing land (coloured dots) on top of a geological map of europe. 97 unfortunately no appropriate horizon 2020 (eu’s framework programme) calls were launched, and two proposals targeting some generic electronic infrastructure calls were rejected. however, the concept of egdi was widely used in strategic communication, and eventually the situation became critical: everybody talked about egdi, but it did not exist. consequently, geus mobilised the so-called ‘spatial information expert group’ of the egs, and after a long process of argumentation and communication the egs general assembly accepted to provide in-kind resources from the national surveys to establish a first basic implementation of the egdi. egdi version 1 on 14 june 2016, the first version of egdi was launched at the premises of the egs secretariat in brussels. besides directors and other egs key persons, the audience comprised a number of distinguished eu commissioners representing different parts of the commission, including dg grow (raw materials), dg mare (maritime affairs), dg ener (geoenergy), dg rtd (research) and dg jrc (joint research and inspire). the participants were very happy with what they saw. even though it is not fully-fledged, egdi v.1 contains most parts of the system that was identified under the egdi-scope project. a number of datasets from past european projects are included and made available through a common web portal (www.europe-geology.eu; fig. 2), which was to a large extent developed by geus. however, the portal is only a small part of egdi. the infrastructure itself consists of a complex of central databases, a metadata catalogue and distributed web services that all conform to the same standards. roles and responsibilities are agreed internally in egs, and work is in progress to establish a more permanent governance structure. egdi in the european landscape of electronic infrastructure egdi is not alone! there are many other related european initiatives (fig. 3) and a few will be mentioned here. the european plate observing system (epos) is a research infrastructure for solid-earth sciences in europe and is of strategic importance to the eu. it aims to facilitate research into natural disasters such as earthquakes and volcanic eruptions through the integration of geoscience data from a number of different research communities. geological data represent one of the domains that epos will integrate; the egs community is the main provider of such data. geus has a role in the implementation project epos-ip and actively uses this position together with other participating national surveys to ensure the complementarity between egdi and epos through various coordination activities. the european marine observation and data network (emodnet) is a large programme aiming at providing uniform access to marine data from europe’s coastal states. emodnet is divided into seven discipline-based themes, one of which deals with geological data. the emodnetgeology project has been running in several phases since it started in 2008. in the current third phase, geus has leading role in the work package dealing with data management and web portal development – again a strong position that geus uses to strengthen the coherence of the european e-infrastructure landscape through technical coordination with egdi, epos and other large initiatives. the role of egdi as a central data-bearing component in european geoscience research is continuously being strengthened through coordination with other european projects and programmes. one of these is geoera. rmis, openecho... fig. 3. the central position of the egdi in the european e-infrastructure landscape. egdi serves geological data from a lot of different geological research areas and thereby contributes geological data to numerous european information systems through interoperable web services. 9898 geoera the first pillar of the egs strategy, the joint research pillar, is on its way to be implemented through an era-net on ‘establishing the european geological surveys research area to deliver a geological service for europe (geoera)’. the main objective of geoera (http://geoera.eu) is to contribute to the optimal use and management of the subsurface. forty-eight national surveys from 33 european countries will participate in research projects under geoera with the purpose of supporting a more integrated and efficient management as well as a more responsible and publicly acceptable exploitation and use of the subsurface. geoera covers the three geoscientific themes of geo-energy, groundwater and raw materials; the projects will run for three years starting in july 2018. the three geoscientific themes share the objective of organising and disseminating a large amount of results in the form of digital maps and other data. furthermore, crossthematic integration of information is an important aspect of geoera. it has therefore been decided to establish a common ‘information platform’ for the efficient integration of all it-related and technical issues (database and dissemination) from all geoera projects. the platform will be capable of integrating up-to-date data, interpretations and models from different and distributed sources, both within and across the three main geoscientific themes of geoera. the platform will be based on egdi which contains some of the functionality foreseen to be required by the geoera projects. egdi will, however, have to be extended in order to handle and disseminate for instance 3d/4d geological models. in addition to organising the geoscientific projects, geoera also carries out ‘foresight activities’ with the purpose of assessing how longer-term funding of the activities can be obtained and eventually lead to a ‘geological service for europe’. the safeguarding of the geoera results through the information platform and the long-term operation and maintenance of the entire egdi should be seen in this context. the future one of the main challenges for all these european initiatives, including egdi, is to make them sustainable. all eu projects are run for a limited period after which the funding ends, and due to budget mechanisms, the european commission is not in a position to fund permanent maintenance of any system. egs has therefore continuously tried to find sustainable funding models to keep egdi alive in a technical and organisational sense. currently the egs members are able to finance the very basic operation of egdi. from 2018–2021 work within the geoera projects will ensure that egdi will be significantly upgraded with much more data and functionality. with egdi becoming a continuously bigger and more integrated piece in the european puzzle, the chances of finding a long-term solution are on the increase, but a sustainable solution will probably have to be connected to the establishment of the ‘geological service for europe’ at least partly funded by eu. other challenges facing the generation of truly paneuropean harmonised geological datasets are that some of the most relevant data are considered confidential in some countries. this is for instance the case for some data about mineral occurrences in certain east european countries. in other countries the organisation and provision of data are the responsibility of regional authorities instead of national surveys, who are less committed to the european viewpoint on data than egs. finally, it has turned out that it is difficult to convert national geological databases and make them interoperable according to the requirements in the inspire implementing rules. many of these rules are very complicated, and many resources have been allocated to the database administrators at the national surveys in order to make their data compliant with the standards. references an effective raw materials strategy for europe. european parliament resolution of 13 september 2011 on an effective raw materials strategy for europe. https://eur-lex.europa.eu (2011/2056(ini)). communication from the commission to the european parliament, the council, the european economic and social committee and the committee of the regions. making raw materials available for europe’s future wellbeing. proposal for a european innovation partnership on raw materials. https://eur-lex.europa.eu (com(2012) 82). kalvig, p. & machacek, e. 2018: examining the rare-earth elements (ree) supply-demand balance for future global wind power scenarios. geological survey of denmark and greenland bulletin 41, 87–90 (this volume). the eurogeosurveys vision towards a geological service for europe. http://www.eurogeosurveys.org/wp-content/uploads/2014/08/egsstrategy-document-2014-a4.pdf. authors’ address geological survey of denmark and greenland, øster voldgade 10, 1350 copenhagen k, denmark. e-mail: jtu@geus.dk. https://eur-lex.europa.eu https://eur-lex.europa.eu http://www.eurogeosurveys.org/wp-content/uploads/2014/08/egs-strategy-document-2014-a4.pdf http://www.eurogeosurveys.org/wp-content/uploads/2014/08/egs-strategy-document-2014-a4.pdf mailto:jtu@geus.dk e2019430103-01 high porosity is a key factor for good reservoir sandstones for both hydrocarbon and geothermal energy exploitation. the porosity of sandstones generally decreases with increased burial depth due to compaction and cementation. however, some sandstones in the north sea show higher porosity than expected for their burial depth, due to the presence of micro­ quartz coatings (e.g. aase et al. 1996; hendry & trewin 1995; jahren & ramm 2000; maast et al. 2011). siliceous sponge spicules have been documented to be an internal source of silica that promotes microquartz coatings (e.g. hendry & trewin 1995; aase et al. 1996). siliceous sponge spicules, the solid ‘skeleton’ of sponges, consist of opal­a and will dissolve when exposed to higher temperatures, thereby causing supersaturation of the formation water with respect to opal­ct and quartz, resulting in nucleation of numerous small (1–5 µm) quartz crystals (williams et al. 1985; hendry & trewin 1995). to predict reservoir quality it is important to understand the distribution of porosity­preserving mic­ roquartz in clastic deposits, and yet this is still poorly un­ derstood. to address this, our study presents petrographical analyses of cored sandstone sections from wells of various depositional environments, including back­barrier, estua­ rine, shoreface and gravity flows, as well as various present­ day burial depths across the danish central graben. geological background the upper jurassic sandstones in the danish central gra­ ben were deposited as part of the syn­rift sedimentation associated with late jurassic rifting of the area. sedimenta­ tion was dominated by mudstones of the lola and farsund formations, with local and episodic deposits of intercalated distribution of porosity-preserving microquartz coatings in sandstones, upper jurassic danish central graben margrethe t. nielsen*1, rikke weibel1, jens therkelsen2 and henrik friis3 geus bulletin is free to individuals and institutions in electronic form. the author(s) retain copyright over the article contents. research article | open access geus bulletin vol 43 | e2019430103 | published online: 22 july 2019 https://doi.org/10.34194/geusb-201943-01-03 jeppe-1 tail end graben salt dome province heno plateau outer rough basin inge high ål basin arne-elin graben feda graben mid north sea high mandal high gert ridge gertrud graben/ plateau coffee soil fault ringkøbing–fyn high søgne basin ravn-1 rita-1 25 km mid north sea high outer moray firth ringkøbing–fyn high central graben uk n d dk 55on m ads h igh ravn-2 w-1 eg-1 diamant-1 gert-4 gwen-2 gert-2 hejre-2 bertel-1 wessel-1 saxo-1 iris-1 tabita-1 lulu-2 nw adda-1 deep adda-1 gert-1 ‘outer rough sand’ heno formation gravity flows ophelia-1ophelia-1 hejre-2 gwen-2 gert-2 gert-1 danish central grabensw ne ringkøbing–fyn high stage ryazanian volgian kimm. oxfordian callovian farsund fm c re ta ce ou s ju ra ss ic se rie s sy st em u pp er lo w er mid north sea high outer rough sand lola fm heno fm ravn mb gert mb cromer knoll lulu fm fig. 1. part of the danish sector of the central gra­ ben showing the 21 sampled wells penetrating upper jurassic sandstones. sandstones with microquartz coatings are indicated by blue dots. modified from johannessen (2003). https://doi.org/10.34194/geusb-201943-01-03 e2019430103-02 sandstones (andsbjerg & dybkjær 2003; johannessen 2003; fig. 1). the heno formation, deposited between the lola formation and the farsund formation, is subdivided into the gert and ravn members, and is intersected locally by the lola formation. the gert member consists of mainly estuarine and back­barrier sandstones, while the ravn mem­ ber comprises mainly shoreface sandstones. both units were deposited in the central part of the danish central graben (fig. 1). the farsund formation sandstones comprise shore­ face sandstones deposited along the mid north sea high (‘outer rough sand’ in fig 1) and deep marine gravity­flow sandstones along the coffee soil fault (fig. 1). since deposi­ tion, the upper jurassic sediments in the danish north sea area have been progressively buried to a maximum depth of 5400 m and temperatures of 160°c (weibel et al. 2019). methods samples were selected from cores of 21 wells penetrating up­ per jurassic sandstones in the danish central graben (fig. 1). plugs were taken perpendicular to the cores, including in one deviated well, and thin sections were made from plug cut­offs. a total of 234 polished thin sections were studied by optical microscopy and by scanning electron microscopy (sem). the relative mineral abundances are assigned as follows: dominant (>50%), abundant (15–50%), common (5–15%), minor (1–5%) and rare (<1%). analyses of crystal 100 µm100 µm 2 µm 50 µm c d a b cq sp mq dq sp sp fig. 2. a: sand laminae rich in sponge spicules (sp) preserved in mud­rich facies (iris­1 well). b: sem­se micrograph showing mouldic porosity from the dissolution of siliceous sponge spicules (arrows) with microquartz coatings (rita­1 well). c: sem­se micrograph showing a detrital quartz grain (dq) coated by cryptocrystalline quartz (cq), and microquartz (mq, rita­1 well). d: optical microscopy image under crossed nicols, showing interlocking quartz overgrowths (arrows, gert­1 well). e2019430103-03 morphologies and paragenetic relationships were performed on gold­coated rock chips by sem. relative chemical com­ positions of mineral phases were obtained using an energy dispersive x­ray spectrometer. plug porosity and perme­ ability were measured according to the api rp­40 standard (api 1998). porosity and permeability measurements were performed on sandstones of the heno formation and from the ‘outer rough sand’ of the farsund formation, only. results detrital composition: in general, quartz grains dominate (>50%) the detrital composition of the upper jurassic sand­ stones, though some samples contain abundant–dominant (i.e. >15%) calcite shell fragments. feldspar, comprising both k­feldspar and albite, varies in abundance from rare (<1%) to common (5–15%). mica rock fragments, organic matter and biogenic particles such as sponge spicules (fig. 2a) and shell fragments are generally rare or minor (1–15%) constituents. heavy minerals such as rutile, zircon, chromite, apatite, fe­ ti­oxides and tourmaline are rare. detrital clays are present in varying amounts. diagenetic changes: early diagenetic (eogenesis) cement in­ cludes pyrite, sporadic anatase, cryptocrystalline and micro­ crystalline quartz coatings. the coatings are occasionally associated with mouldic porosity – a form of secondary po­ rosity formed by the dissolution of siliceous sponge spicules. we also observed chlorite coatings in samples with detrital volcanic particles, pores filled by calcite cement in samples with abundant detrital carbonate particles and secondary porosity from the dissolution of feldspar and kaolinite. later diagenetic (mesogenesis) phases include illite, rare feldspar overgrowths, and dolomite and ankerite overgrowths on detrital carbonate clasts. also observed were quartz over­ growths, late­patchy calcite cement, rare siderite, barite and sphalerite. siliceous authigenic phases: microcrystalline quartz coat­ ings are observed in a number of wells (fig. 1). microquartz typically occurs as a coating on detrital quartz grains (figs. 2c, d), and as a coating on detrital carbonate grains in sand­ stones from the nw adda­1 well. cryptocrystalline quartz is observed between the detrital grains and the microquartz coating as shown in fig. 2c. sponge spicules, replaced by the quartz variant chalcedony, and mouldic porosity after dis­ solved sponge spicules are occasionally present in samples with microquartz coatings (figs 2a, b). microcrystalline silica is very abundant and has lithified the sandstones from the tabita­1 and deep adda­1 wells, which contain abun­ dant siliceous sponge spicules. quartz overgrowths vary in size from discrete incipient overgrowths, which may have started as small outgrowths, to large interlocking quartz ce­ mentation as in sandstones from the gert­1 well (fig. 2d). distribution of microquartz: stratigraphically, micro­ quartz coatings are restricted to the farsund formation and ravn member of the heno formation, but absent in the lola formation and in gert member of the heno formation (fig. 0 20 40 60 80 100 120 back-barrier estuarine tidal upper shoreface middle shoreface lower shoreface shoreface offshore gravity flow d ep os iti on al en vir on m en t number of samples 0 20 40 60 80 100 120 ravn mb gert mb lola fm farsund fmli th os tr at igr ap hy with microquartz no microquartz a b fig. 3. distribution of microquartz according to a: lithostratigraphy and b: depositional environment. here, offshore environments are represented by the lola formation. e2019430103-04 3). microquartz coatings are restricted to shoreface and grav­ ity­flow sandstones, whereas microquartz coatings are absent in sandstones deposited in back­barrier, estuarine and tidal environments (fig. 3). porosity variation: sandstones of the heno formation generally plot with relatively high porosities when compared to the regional mean porosity­depth trends given by selley (1978) and bjørkum et al. (1998) (fig. 4). samples with micro quartz, shallower than 4.6 km, have overall porosi­ ties of 20.7% ± 8.0, compared with 12.2% ± 5.6 for samples without microquartz. many samples deeper than 4.6 km also show relatively high porosities compared to regional trends, even though microquartz is rarely observed in these sand­ stone samples (weibel et al. 2019). discussion the porosity of the upper jurassic sandstones shows large variations at all depths (fig. 4), though there is a general trend towards lower average porosity at increased depth. microquartz has a positive influence on porosity preserva­ tion in sandstones from depths ≤4.6 km, which is the maxi­ mum depth of observed microquartz in the danish central garben. this is due to a combination of mechanical stabili­ sation of the sandstone and inhibition of quartz cementation (aase et al. 1996; jahren & ramm 2000; bonnell et al. 2006; lander et al. 2008). the porosity­preserving effect of mic­ roquartz coatings is particularly pronounced when the indi­ vidual crystals are randomly orientated. whereas coatings with a consistent crystallographic orientation can become overgrown by macroquartz (e.g. weibel et al. 2010; french et al. 2012; french & worden 2013). the presence of microquartz coatings surrounding moul­ dic porosity after dissolved sponge spicules (fig. 2b) suggests a close relationship, which is in line with previous studies by hendry & trewin (1995) and maast et al. (2011), suggest­ ing that microquartz coatings in the north sea are sourced from dissolved silicous sponge spicules. in the upper juras­ sic sandstones of the danish central graben, microquartz coatings are restricted to sandstones deposited in gravity­ flow and shoreface environments (ravn member and far­ sund formation), while microquartz is absent in sandstones deposited in back­barrier and estuarine environments (gert member). this corresponds to the preferred marine habi­ tats of silica sponges (demosponges) during the late jurassic (leinfelder et al. 1996). demosponges are filter feeding or­ ganisms that consume predominantly nanoplankton, which decrease in abundance with depth, and hence, demosponges preferentially occur in shallow marine, middle and outer ramp environments (leinfelder et al. 1996). as the sponges colonised the sea floor, died and disintegrated, the ‘skeletal’ spicules, would have been released into the environment. sponge spicules from the shoreface environments could have been carried to deeper waters along with sand grains by gravity flow, similar to the interpretations of the cretaceous sandstones in the uk north sea (hendry & trewin 1995). microquartz coatings are only observed in sandstones buried no deeper than than 4.6 km (fig. 4), corresponding to maximum temperatures of c. 150oc assuming a geothermal gradient of 33oc km–1 (evans 2003). this may be explained by increasing temperatures making larger quartz crystals in continuous overgrowths thermodynamically more stable (williams et al. 1985). microcrystalline quartz formed at a low temperature may coarsen during increased burial to form a thermodynamicaly stable crystal, but this thermodynamic drive is less likely to have promoted coarse microquartz crys­ tals (c. 5–10 µm) to be dissolved and replaced by syntaxial overgrowths (william et al. 1985; hendry & trewin 1995; jahren & ramm 2000). the deeper buried sandstones prob­ ably never developed significant amounts of microquartz or other early diagenetic phases, such as opal­ct, or zeolite protected the grain surfaces and were later dissolved during increased burial, thereby exposing the grain surfaces for mac­ roquartz nucleation (e.g. hendry & trewin 1995; weibel et al. 2010, 2019). 2500 3000 3500 4000 4500 5000 5500 6000 d ep th  (m ) porosity (%) 0 5 10 15 20 25 30 35 samples without microquartz samples with microquartz average porosity at  200 m intervals expected porosity (selley 1978) expected porosity (bjørkum et al. 1998) fig. 4. porosity vs. depth for samples with (white dots) and without (black bots) microquartz. also shown are estimated or predicted porosity­depth curves according to selley (1978) and bjørkum et al. (1998). e2019430103-05 conclusions microquartz coatings in upper jurassic sandstones of the danish central graben are confined to offshore gravity flows and shoreface sandstones, and absent in sandstones deposited in back­barrier and estuarine environments. in the samples presented here, chalcedonic sponge spicules and mouldic porosity after sponge spicules often coincide with the presence of microquartz coatings. this observation supports the idea that the presence of microquartz is deter­ mined by silica sponge habitats and depositional processes. improved understanding of the depositional environments likely to concentrate siliceous sponge spicules would enable prediction of sandstone reservoirs that contain microquartz coatings and exceptionally high porosity and permability. acknowledgments this study is based on results from the multiclient project petsys to which numerous colleagues at geus have contributed. thanks are ad­ dressed to reviewers j. hendry and j. jahren for their consructive com­ ments, which improved the paper. references aase, n.e., bjorkum, p.a. & nadeau, p.h. 1996: the effect of grain­ coating microquartz on preservation of reservoir porosity. aapg bul­ letin 80, 1654–1673. https://doi.org/10.1306/64eda0f0­1724­11d7­ 8645000102c1865d andsbjerg, j. & dybkjær, k. 2003: sequence stratigraphy of the jurassic of the danish central graben. geological survey of denmark and green­ land bulletin 1, 265–300. api rp 40. 1998: recommended practices for core analysis, second edi­ tion. 236 pp. washington: american petroleum institute. bjørkum, p.a., oelkers, e.h., nadeau, p.h., walderhaug, o. & murphy, w.m. 1998: porosity prediction in quartzose sandstones as a function of time, temperature, depth, stylolite frequency, and hydrocarbon satura­ tion. aapg bulletin 82, 637–648. https://doi.org/10.1306/1d9bc5cf­ 172d­11d7­8645000102c1865d bonnell, l., larese, r. & lander, r. 2006: porosity preservation by inhibi­ tion of quartz cementation: microquartz versus hydrocarbons. aapg international conference and exhibition, perth, west australia, 5–6 november, 2006. evans, d. 2003: the millennium atlas: petroleum geology of the central and northern north sea, 389 pp. london: the geological society. french, m.w. & worden, r. 2013: orientation of microcrystalline quartx in the fontainebleau formation, pris basin and why it pre­ serves porosity. sedimentary geology 284–285, 149–158. https://doi. org/10.1016/j.sedgeo.2012.12.004 french, m.w., worden, r., mariani, e., larese, r.e., mueller, r.r. & kliewer, c.e. 2012. microcrystalline quartz generation and the pres­ ervation of porosity in sandstones: evidence from the upper cretaceous of the subhercynian basin, germany. journal of sedimentary research 82, 422–434. https://doi.org/10.2110/jsr.2012.39 hendry, j.p. & trewi, n. 1995: authigenic quartz microfabrics in creta­ ceous turbidites: evidence for silica transformation processes in sand­ stones. journal of sedimentary research a65, 380–392. https://doi. org/10.1306/d42680cc­2b26­11d7­8648000102c1865d jahren, j. & ramm, m. 2000: the porosity­preserving effects of micro­ crystalline quartz coatings in arenitic sandstones: examples from the norwegian continental shelf. in: worden, r.h. & morad, s. (eds): quartz cementation in sandstones: special publication 29 of the in­ ternational association of sedimentologists series, 271–280. blackwell publishing ltd. https://doi.org/10.1002/9781444304237.ch18 johannessen, p.n. 2003: sedimentology and sequence stratigraphy of paralic and shallow marine upper jurassic sandstones in the northern danish central graben. geological survey of denmark and greenland bulletin 1, 367–402. lander, r.h., larese, r.e. & bonnell, l.m. 2008: toward more accu­ rate quartz cement models: the importance of euhedral versus non­ euhedral growth rates. aapg bulletin 92, 1537–1563. https://doi. org/10.1306/07160808037 leinfelder, r.r., werner, w., nose, m., schmid, d.u., krautter, m., lat­ ernser, r., takacs, m. & hartman, d. 1996: paleoecology, growth pa­ rameters and dynamics of coral, spone and microbolite reefs from the late jurassic. in: reitner, j., neuweiler, f. & gunkel, f. (eds): global and regional controls on biogeneic sedimentation. i. reef evolution. research reports. sb2, 227–248. universität stuttgart. maast, t.e., jahren, j. & bjørlykke, k. 2011: diagenetic controls on res­ ervoir quality in middle to upper jurassic sandstones in the south vi­ king graben, north sea. aapg bulletin 95, 1937–1958. https://doi. org/10.1306/03071110122 selley, r.c. 1978: porosity gradients in north sea oil­bearing sand­ stones. journal of the geological society 135, 119–132. https://doi. org/10.1144/gsjgs.135.1.0119 weibel, r., friis, h., kazerouni, a.m., svendsen, j.b., stokkendahl, j. & poulsen, m.l. 2010: development of early diagenetic silica and quartz morphologies – examples from the siri canyon, danish north sea. sedimentary geology 228, 151–170. https://doi.org/10.1016/j.sedg­ eo.2010.04.008 weibel, r., olivarius, m., jakobsen, f.c., whitehouse, m., midtgaard, h., larsen, m. & nielsen, k. 2019: thermogenetic degradation of early zeolite cement: an important process for generating anomalously high porosity and permeability in deeply buried sandstone reservoirs? ma­ rine and petroleum geology 103, 620–645. https://doi.org/10.1016/j. marpetgeo.2019.02.006 williams, l.a., parks, g.a. & crerar, d.a. 1985. silica diagenesis: i. solu­ bility controls. journal of sedimentary petrology 55, 301–311. https:// doi.org/10.1016/0198­0254(85)93828­2 how to cite nielsen, m.t., weibel, r., therkelsen, j. & friis, h. 2019: distribu­ tion of porosity­preserving microquartz coatings in sandstones, up­ per jurassic danish central graben. geological survey of denmark and greenland bulletin 43, e2019430103. https://doi.org/10.34194/ geusb-201943-01-03 *corresponding author: margrethe t. nielsen | e-mail: mtn@geus.dk 1 geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350, copenhagen k, denmark. 2moe a/s, næstvedvej 1, dk-4760 vordingborg, denmark 3department of geoscience, aarhus university, høegh-guldbergs gade 2, dk-8000 aarhus c, denmark. https://doi.org/10.1306/64eda0f0-1724-11d7-8645000102c1865d https://doi.org/10.1306/64eda0f0-1724-11d7-8645000102c1865d https://doi.org/10.1306/1d9bc5cf-172d-11d7-8645000102c1865d https://doi.org/10.1306/1d9bc5cf-172d-11d7-8645000102c1865d https://doi.org/10.1016/j.sedgeo.2012.12.004 https://doi.org/10.1016/j.sedgeo.2012.12.004 https://doi.org/10.2110/jsr.2012.39 https://doi.org/10.1306/d42680cc-2b26-11d7-8648000102c1865d https://doi.org/10.1306/d42680cc-2b26-11d7-8648000102c1865d https://doi.org/10.1002/9781444304237.ch18 https://doi.org/10.1306/07160808037 https://doi.org/10.1306/07160808037 https://doi.org/10.1306/03071110122 https://doi.org/10.1306/03071110122 https://doi.org/10.1144/gsjgs.135.1.0119 https://doi.org/10.1144/gsjgs.135.1.0119 https://doi.org/10.1016/j.sedgeo.2010.04.008 https://doi.org/10.1016/j.sedgeo.2010.04.008 https://doi.org/10.1016/j.marpetgeo.2019.02.006 https://doi.org/10.1016/j.marpetgeo.2019.02.006 https://doi.org/10.1016/0198-0254(85)93828-2 https://doi.org/10.1016/0198-0254(85)93828-2 https://doi.org/10.34194/geusb-201943-01-03 https://doi.org/10.34194/geusb-201943-01-03 mailto:mtn%40geus.dk?subject= geological survey of denmark and greenland bulletin 38, 2017, 61-64 61 the sedimentary record of the glaciated margins of the north atlantic holds evidence of past ice-sheet activity, and reflects spatial and temporal variations in the ice– ocean–climate interaction as well as the influence of tectonic processes. furthermore, the record of cross-shelf ice sheets provides a direct link between the continental ice cover and the deep ocean, a relevant issue in the context of climate research. with a four-year funding period, a marie curie initial training network on the glaciated north atlantic margins (glanam) was started in the spring 2013. the network involved international partners from both academia and industry and enrolled 15 young scientists working in different areas of the north atlantic margins. the geological survey of denmark and greenland was partner in the network, leading the greenland margin research and hosting three fellows. one of the main topics of the glanam project was to investigate the impact of the ice sheets on the large-scale evolution of the east greenland margin. the present work summarises some new insights into the glacial history of the central east greenland margin gained through a study within this project (fig. 1). the glaciated east greenland margin the dynamic evolution of the greenland ice sheet is related to the glacial history of the northern hemisphere (e.g. thiede et al. 2010). despite evidence of glaciated hinterland and tidewater glaciers in greenland during the eocene and miocene (e.g. thiede et al. 2010), the onset of large-scale glaciations in the northern hemisphere, where ice expanded onto the continental shelf, has been suggested to date between 5 and 2.5 ma based on a marked decrease in global benthic δ18o values and the presence of ice-rafted debris deposits (e.g. bailey et al. 2013). since the midpliocene the oscillation of the ice sheets of the northern hemisphere, and thus of the east greenland margin, is considered to have followed the glacial–interglacial cycles (e.g. sarnthein et al. 2009). major tectonic events related to the miocene–pliocene uplift of the east greenland margin have been pointed to as instigators of the eastwards glacial advance across the shelf (døssing et al. 2016). furthermore, the build-up of the north atlantic ice sheets has also been influenced by the oceanographic circulation which, along east greenland, is mainly controlled by the east greenland current (egc; fig. 1; sarnthein et al. 2009). the egc is a southward-flowing current formed by a complex system of branches and different water masses (våge et al. 2013). asynchronous ice-sheet development along the central east greenland margin: a glanam project contribution lara f. pérez and tove nielsen 15°w denmark strait kong oscar fjord kejser franz joseph fjord kangerlussuaq scoresby sund ce nt ra l e as t g re en lan d m ar gin geikie plateau blo sse vil le ky st iceland 25°w30°w 20°w 67°n 68°n 73°n egc egc jameson land geikie ll fig. 1. bathymetric map of the study area based on the international bathymetric chart of the arctic ocean (ibcao; jacobsson et al. 2012). contour interval: 500 m. orange lines: reflection seismic profiles. red lines: sub-bottom and multibeam data. red dot: odp site 987. black arrows: general circulation of the east greenland current (egc). ll: liverpool land. note the outwards bulging of the shelf edge off the major fjords. © 2017 geus. geological survey of denmark and greenland bulletin 38, 61–64. open access: www.geus.dk/publications/bull 6262 it is an important component of the atlantic meridional overturning circulation (amoc) and therefore has implications for the global climate system (de schepper et al. 2015). the contribution of the egc to the amoc has been partly controlled by tectonic pulses of the denmark strait (parnell-turner et al. 2015), a 600 m deep threshold located around 67°n (fig. 1). the denmark strait is part of the greenland–scotland ridge and constitutes a natural boundary that divides the east greenland margin into a northern and southern part, which have experienced different ice-sheet dynamics. our study focuses on the glacial history of the margin section that lies just north of the denmark strait, i.e. the central east greenland margin (fig. 1). database and interpretation procedure based on a large database of 2d seismic reflection profiles (fig. 1), we have divided the sedimentary record into major seismic sequences, which show evidence of various sedimentary processes (pérez et al. unpublished data). a local dataset of high-resolution reflection seismic, sub-bottom profiles and swath bathymetry allowed for a detailed breakdown of the recent stratigraphic and morphological features (pérez et al. unpublished data). the ages of the major seismic sequences have been assessed by a direct tie to site 987 of the ocean drilling program off scoresby sund (fig. 1; pérez et al. unpublished data ). seismic-stratigraphic analyses of these datasets made a large-scale reconstruction of key stratigraphic events possible, revealing several stages of the greenland ice sheet dynamics along the central east greenland margin from late miocene to present. evidence for cross-shelf ice sheets and ice streams central east greenland is characterised by large fjords, many of which are connected with cross-shelf troughs that are up to 300 m deep and 35 km wide (fig. 1). these troughs were formed by erosion of ice streams that passed from the fjords across the shelf, delivering a concentrated accumulation of sediments to the shelf edge (e.g. batchelor & dowdeswell 2014). from there the sediments were transported down the slope as glacigenic debris flows that build up to form large prograding wedges called troughmouth fans. the presence of a trough-mouth fan is often revealed in the seabed morphology as an outwards bulging of the shelf edge (fig. 1), and the youngest debris flows are often observed on the present-day seabed. ice streams are recognised as one of the most important controls on ice-sheet configuration and stability (e.g. stokes et al. 2016). therefore the study of palaeo-ice stream behaviour and dynamics by mapping buried cross-shelf troughs and trough-mouth fans is a useful tool in reconstructing former ice-extent and palaeoclimate variability (e.g. batchelor & dowdeswell 2014). in addition, submarine glacial forms outside the cross-shelf troughs hold clues of the existence of more steady, grounded ice. notable features are grounding-zone wedges up to 160 m high identified along the central east greenland shelf. these ridge-like sedimentary features mark a temporary position of the ice margin on the shelf (dowdeswell & fugelli 2012). discussion prograding deposits off blosseville kyst dating back to the late miocene constitute the first evidence of cross-shelf glaciations on the central east greenland margin (fig. 2). the oceanward glacial advance continued during the early pliocene, where cross-shelf troughs and trough-mouth fans off blosseville kyst and scoresby sund denote ice-sheet growth with ice streams occasionally reaching the palaeo shelf edge (fig. 2). this glacial intensification coincided with the first large-scale glaciation reaching to the palaeoshelf edge along the south-western greenland margin (nielsen & kuijpers 2013). during the middle pliocene (3.65–2.90 ma), the seismic-stratigraphic analysis denotes a period of glacial retreat along the central east greenland margin. as the ice retreated, the oceanic current took over the control of the depositional environment, indicated by a predominance of current-generated wavy facies. the observed glacial retreat is coeval with the global mid-pliocene warmth (3.3–3.0 ma; e.g. robinson 2009) and a supposed enhancement of the egc along the east greenland margin (e.g. raymo et al. 1996). thick trough-mouth fan deposits led to a major oceanward advance of the shelf edge off scoresby sund, providing evidence of multiple cross-shelf glaciations during the quaternary (fig. 2). the ice-sheet extension was largest during latest pliocene – earliest pleistocene (2.90–2.33 ma), revealing a slightly older age for the onset of margin progradation off central east greenland than observed farther north off north-east greenland (c. 76°n), where the first margin progradation began c. 2.5 ma (berger & jokat 2009). the large progradation of the central east greenland margin coincided with the proposed onset of major northern hemisphere cooling at 2.7 ma (e.g. bailey et al. 2013) and the suggested full-scale glaciation of greenland at 2.9 ma (sarnthein et al. 2009). 63 in addition to the cross-shelf troughs off blosseville kyst and scoresby sund, grounding-zone wedges are identified on the shelf off liverpool land within the quaternary sequences, providing evidence of steady, grounded ice. thus, repeated glacial advances over the shelf, occasionally reaching the shelf edge, are inferred along the entire central east greenland margin during the quaternary (fig. 2). however, the study of glacigenic debris-flow deposits observed on the high-resolution dataset of the liverpool land margin (fig. 1) indicates that the quaternary glacial advances to the shelf edge were not synchronous along the margin. the glacigenic debris-flow deposits identified within the early pleistocene sequences in the southern part of the liverpool land dataset suggest a distal downslope input from the scoresby sund ice stream, in agreement with higher sediment supply to the north of the scoresby sund trough-mouth fan between 1.77 and 0.78 ma (laberg et al. 2013). an upward increase of glacigenic debris-flow deposits within the upper seismic section indicates an intensification of glacial control on the sedimentation during the middle pleistocene. this scenario matches the increase in global ice volume that accompanied the mid-pleistocene transition c. 0.9–0.8 ma (head & gibbard 2005) and gave rise to the growth of larger ice sheets in the northern hemisphere (e.g. dowdeswell et al. 1997; stokes et al. 2016). the internal distribution of the middle pleistocene glacigenic debris-flow deposits points to a changing sediment source through time. whereas the oldest glacigenic debris-flow deposits are most abundant in the southern part of the liverpool land area, pointing to an ice-stream source in scoresby sund, the youngest glacigenic debrisflow deposits are more abundant in the northern part of the study area and thus are likely feed by an ice stream from kong oscar fjord (fig. 2). this northern-sourced pattern continued during the latest pleistocene and holocene, in agreement with the presence of ice-rafted debris trapped inside scoresby sund during the last 10 ka (stein et al. 1993) and the southward-pointing, cross-shelf trough off this fjord observed in the present-day seafloor (dowdeswell et al. 1997). farther north, moraines related to the maximum extent of the greenland ice sheet during the last glacial maximum have been identified off kejser franz joseph fjord (evans et al. 2002). shelf slope blosseville kyst liverpool l shelf slope blosseville kyst liverpool l eg c eg c tmf gdf t tmf gdf t ss ss ka ka quaternary late miocene – early pliocene ko ko basin plain basin plain glacial erosion fig. 2. 3d sketch of the central east greenland margin development during the late miocene – early pliocene and the quaternary showing the main morphological features and key depositional processes. within the same age range, darker colours represent older processes or deposits. t: trough. tmf: trough-mouth fan. gdf: glacigenic debris-flow deposit. blue lines on shelf: predominantly glacial erosion. orange lines on slope: predominantly progradation. blue arrows: egc: east greenland current. ka: kangerlussuaq. ss: scoresby sund. ko: kong oscar fjord. 6464 concluding remarks our data indicate an early cross-shelf glaciation off blosseville kyst during the late miocene and early pliocene followed by major cross-shelf glaciations off scoresby sund during the early quaternary and off liverpool land in the late quaternary. higher resolution of the quaternary data off liverpool land indicates that the activity of the scoresby sund ice-stream system was gradually taken over by the kong oscar fjord ice-stream system during the pleistocene. overall, our study reveals an asynchronous growth of the ice sheet across the shelf, with a marked northward progradation of ice-stream activity from the late miocene to the present along the central east greenland margin. acknowledgements the research leading to these results received funding from the people programme (marie curie actions) of the eu fp7 programme fp7/2007-2013/ under rea grant agreement no. 317217. the research forms part of the glanam (glaciated north atlantic margins) initial training network. for further information on the glanam project visit www.glanam.org. references bailey, i., hole, g.m., foster, g.l., wilson, p.a., storey, c.d., trueman, c.n. & raymo, m.e. 2013: an alternative suggestion for the pliocene onset of major northern hemisphere glaciation based on the geochemical provenance of north atlantic ocean ice-rafted debris. quaternary science reviews 75, 181–194. batchelor, c.l. & dowdeswell, j.a. 2014: the physiography of high arctic cross-shelf troughs. quaternary science reviews 92, 68–96. berger, d. & jokat, w. 2009: sediment deposition in the northern basins of the north atlantic and characteristic variations in shelf sedimentation along the east greenland margin. marine and petroleum geology 26, 1321–1337. de schepper, s., schreck, m., beck, k.m., matthiessen, j., fahl, k. & mangerud, g. 2015: early pliocene onset of modern nordic seas circulation related to ocean gateway changes. nature communications 6, 8659, http://dx.doi.org/10.1038/ncomms9659 døssing, a., japsen, p., watts, a.b., nielsen, t., jokat, w., thybo, h. & dahl–jensen, t. 2016: miocene uplift of the ne greenland margin linked to plate tectonics: seismic evidence from the greenland fracture zone, ne atlantic. tectonics 35, 257–282. dowdeswell, j.a. & fugelli, e.m.g. 2012: the seismic architecture and geometry of grounding-zone wedges formed at the marine margins of past ice sheets. geological society of america bulletin 124, 1750–1761. dowdeswell, j.a., kenyon, n.h. & laberg, j.s. 1997: the glacierinfluenced scoresby sund fan, east greenland continental margin: evidence from gloria and 3.5 khz records. marine geology 143, 207–221. evans, j., dowdeswell, j.a., grobe, h., niessen, f., stein, r., hubberten, h.w. & whittington, r.j. 2002: late quaternary sedimentation in kejser franz joseph fjord and the continental margin of east greenland. geological society special publications (london) 203, 149–179. head, m.j. & gibbard, p.l. 2005: early-middle pleistocene transitions: an overview and recommendation for the defining boundary. geological society special publications (london) 247, 1–18. jakobsson, m. et al. 2012: the international bathymetric chart of the arctic ocean (ibcao) version 3.0. geophysical research letters 39, l12609. laberg, j.s., forwick, m., husum, k. & nielsen, t. 2013: a re-evaluation of the pleistocene behavior of the scoresby sund sector of the greenland ice sheet. geology 41, 1231–1234. nielsen, t. & kuijpers, a. 2013: only 5 southern greenland shelf edge glaciations since the early pliocene. scientific reports 3, 1875. parnell-turner, r., white, n.j., mccave, i.n., henstock, t.j., murton, b. & jones, s.m. 2015: architecture of north atlantic contourite drifts modified by transient circulation of the icelandic mantle plume. geochemistry, geophysics, geosystems 16, 3414–3435. raymo, m.e., grant, b., horowitz, m. & rau, g.h. 1996: mid-pliocene warmth: stronger greenhouse and stronger conveyor. marine micropaleontology 27, 313–326. robinson, m.m. 2009: new quantitative evidence of extreme warmth in the pliocene arctic. stratigraphy 6, 265–276. sarnthein, m., bartoli, g., prange, m., schmittner, a., schneider, b., weinelt, m., andersen, n. & garbe-schönberg, d. 2009: mid-pliocene shifts in ocean overturning circulation and the onset of quaternary-style climates. climate of the past discussions 5, 269–283. stein, r., grobe, h., hubberten, h., marienfeld, p. & nam, s. 1993: latest pleistocene to holocene changes in glaciomarine sedimentation in scoresby sund and along the adjacent east greenland continental margin: preliminary results. geo-marine letters 13, 9–16. stokes, c.r., margold, m., clark, c.d. & tarasov, l. 2016: ice stream activity scaled to ice sheet volume during laurentide ice sheet deglaciation. nature 530, 322–326. thiede, j., jessen, c., knutz, p., kuijpers, a., mikkelsen, n., nørgaardpedersen, n. & spielhagen, r.f. 2010: millions of years of greenland ice sheet history recorded in ocean sediments. polarforschung 80, 141–159. våge, k., pickart, r.s., spall, m.a., moore, g.w.k., valdimarsson, h., torres, d.j., erofeeva, s.y. & nilsen, j.e.ø. 2013: revised circulation scheme north of the denmark strait. deep sea research part i. oceanographic research papers 79, 20–39. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark; e-mail: lfp@geus.dk geological survey of denmark and greenland bulletin 38, 2017, 9-12 9 there is a growing demand in modern society for detailed, localised geological maps and 3d models in connection with e.g. planning of major construction works, study of subsurface drainage systems, infiltration of storm water or risk assessment of contaminated waste dumps and pollution plumes. this demand is difficult to meet in denmark as the surficial glacial deposits that cover most of the country are notoriously very heterogeneous. standard geological maps are based on regional data collection, and their resolution is far from sufficient to identify structural elements on the 10–20 m scale needed in the above-mentioned applications. geophysical mapping for geological characterisation of the upper c. 5 m of the subsurface can be carried out using for instance direct-current geoelectrical methods (e.g. loke et al. 2013), induced polarisation (e.g. revil et al. 2012) set up with 1–2 m electrode spacing, electromagnetic induction (emi; e.g. christiansen et al. 2016; doolittle & brevik 2014), ground penetrating radar (gpr; e.g. neal 2004) or seismic refraction tomography using a multicomponent landstreamer (e.g. brodic et al. 2015). the resulting geophysical maps show the distribution of the measured parameter, for instance electric resistivity or seismic velocity. to construct geological maps using geophysical methods, the data must be verified and calibrated with geological field observations. gpr imaging of geological structures require laborious interpretation before a geological map can be constructed, and the method is limited to low-loss materials such as sandy sediments (neal 2004). a new approach, using a combination of shallow, highresolution emi surveying and traditional spear-auger soil sampling along the same transects, was tested in an area of c. 2 km2 around the contaminated, former landfill site at pillemark on samsø (fig. 1). the resistivity recorded using the emi method is strongly related to the clay content, and this parameter is therefore well suited for geological mapping. the emi method is also robust, data acquisition is 5–50 times faster than with other geophysical methods and the processing and inversion scheme is well defined (christiansen et al. 2016). methodology spear-auger mapping in the past almost 130 years, the geological mapping of the surficial cover of denmark has largely been based on simple collection of pristine samples of the local sediment below the mull horizon using a specially designed sampling device, the so-called spear auger. this is a 1 m long steel rod with a diameter of 12 mm, a handle bar, and a 15–25 cm long, single or double slit at its tip. when the spear is pushed into the ground and turned, the slit captures a small soil sample (fig. 2). the mapping geologist interprets the nature and origin of the sample and adds a soil-type symbol to a field map currently using the terminology described by jakobsen et al. (2011). to map the boundaries between different soil types, samples are collected with a distance of 100–200 m depending on the local geological complexity. the symbols are then transferred to a master optimising geological mapping of glacial deposits using high-resolution electromagnetic induction data knud erik s. klint, ingelise møller, pradip k. maurya, and anders v. christiansen elevation relative to sea level 50 m 25 m 0 m -25 m -50 m denmark 100 km 2 km fig. 1. the location of the study area on the island of samsø is shown on the digital terrain model. © 2017 geus. geological survey of denmark and greenland bulletin 38, 9–12. open access: www.geus.dk/publications/bull 1010 map sheet on a scale of 1:25 000; the resolution of the resulting geological map is about ±100 m. geophysical mapping by electromagnetic induction emi methods are commonly used for soil mapping (e.g. doolittle & brevik 2014). during the last decade the development of multi-coil sensors, integration with gps and use of inversion algorithms for data interpretation (e.g. christiansen et al. 2016; doolittle & brevik 2014), have made emi a very fast technique for mapping the resistivity of the upper 4–8 m of the subsurface (fig. 3). the multicoil dualem-421s sensor used in our study transmits an electromagnetic wave at 9 khz from a horizontal coil and samples the total signal in horizontal and vertical receiver coils 1, 2 and 4 m from the transmitter coil. a signal is received 10 × per second, making the equipment suitable for towing behind a motorised vehicle (fig. 3). acquisition of field data in the study area the emi survey of the study area comprised c. 85 linekm mainly by towing; in difficult terrain and vegetation the sensor was carried manually. the emi data were (1) averaged using a running mean filter of 2 m, (2) outliers manually culled and (3) inverted using a full-solution 1d algorithm (auken et al. 2015), where the models are linked by 3d constraints to obtain a pseudo-3d resistivity model (viezzoli et al. 2008). the resistivity models were discretised in 10 layers covering the upper 10 m of the subsurface; the mean depth of investigation was c. 6 m (christiansen & auken 2012). the interval resistivity at 1–2 m below surface was calculated (fig. 4) and used together with the spear-auger mapping. the emi survey was conducted prior to the spear-auger mapping, and the selection of soil sampling points was adjusted to the interval resistivity map so that areas with highly variable resistivity were mapped at a resolution of 10–20 m between the sampling points and areas with more homogeneous resistivity were mapped at a lower resolution. in fig. 4 soil symbols and transitions of polygons with similar soil types based primarily on the spear-auger mapping are shown as an overlay on the geophysical mapping. the final interpretation of the soil type distribution is shown in fig. 5, where it is seen that the soil type boundaries could be drawn with much higher accuracy than with traditional spear-auger mapping. results the pillemark study area (fig. 1) has a hummocky topography. the geological map (fig. 5) shows that a large variety of soil types is present. clayey deposits such as clay till partly overlying meltwater clay dominates the hills, whereas sandier deposits such as sandy tills, meltwater sand/gravel occur along the fringes of the hills and on the highest hilltops. extramarginal sand predominates in the low-lying areas, in the lowermost areas and in local depressions partly covered with postglacial sediments such as freshwater sand, clay and peat. based on the combined mapping techniques, the depositional history of the soil types could be outlined. the hills are interpreted to consist of mainly layered lacustrine deposits, which were overridden by a glacier that deposited tills and potentially deformed the lake sediments before fig. 2. traditional geological mapping using the spear auger. fig. 3. the emi survey system towed behind a vehicle. the c. 4 m long sensor is located inside a white tube on two sleds. the gps sensor is mounted in the front of the tube above the transmitter coil. a data logger and a computer controlling the data acquisition are placed on the vehicle. 11 pillemark n ?? ? ??????? a b late glacial deposits ts glacial deposits ml clay till ml sandy clay till ml gravelly clay till ms sandy till dg meltwater gravel ds meltwater coarse sand dl meltwater clay postglacial deposits fg freshwater gravel fs freshwater sand fl freshwater clay/silt ft freshwater peat fp freshwater gytja extramarginal freshwater sand late glacial deposits ts extramarginal freshwater sand glacial deposits dg meltwater gravel ds meltwater sand dl meltwater clay ms sandy till ml clayey till postglacial deposits fs freshwater sand fl freshwater clay ft freshwater peat fill/town pillemark landfill site 200 m 200 m 100 1000 ohm/m 10 35 30 25 500 m100 200 300 400 profilea bm a.s.l. fig. 4. raw data from the spear-auger mapping and the interval resistivity at 1–2 m depth in the survey area. note how the boundaries between soil types as interpreted in the field on this figure were modified during the construction of the geological map (fig. 5). note also the strong dark blue signature of the pillemark landfill site, indicating the distribution of contaminated soil. fig. 5. high-resolution geological map of the research area based on combined spear-auger mapping and emi surveying. 1212 stagnating and melting. lakes in local depressions were then slowly filled with postglacial freshwater sand, silt/clay and organic matter, slowly transforming into peat bogs. the landfill site was covered with anthropogenic soil and fill material; the outline of the contaminated soil is clearly apparent in fig. 4. summary and perspectives compared to traditional spear-auger mapping, the combined emi and spear-auger mapping makes it possible to produce high-resolution geological maps in highly heterogeneous glacial landscapes. it efficiently outlined the contaminated area at the landfill site. the method is considered solid and very cost-effective, since the total mapping of the area was carried out in three days plus another three days’ work for processing the geophysical data and correlation with the spear-auger mapping. unfortunately, local disturbance from buried pipes, electrical wires etc. influences the quality of the geophysical data within a distance of 10–15 m from the instrument, making the technique less applicable in densely urbanised areas. moreover, without using a gps, the precision of spear-auger mapping is low on a 10-metre scale. a dedicated mapping exercise coordinating gps positioning of all sampling points with geophysical measurements would increase the overall resolution. finally, spear-auger mapping requires a well-trained geologist for interpreting the very small soil samples correctly. the combined emi and spear-auger mapping also has great potential for site characterisation, from general surveys on a scale of c. 10 km to local studies on a scale of c. 100 m. the methodology can easily be adapted to, or combined with, more dedicated soil sampling e.g. for detection of contaminated soil. there is also an obvious potential in other situations where high-resolution geological maps are needed, for example for identification of efficient infiltration areas in connection with establishment of suburban drainage systems (bockhorn et al. 2015). acknowledgements this study was supported by the research project geocon – advancing geological, geophysical and contaminant monitoring technologies for contaminated site investigation (contract 1305-00004b). funding for this is provided by the danish council for strategic research under the programme commission on sustainable energy and environment. references auken, e. et al. 2015: an overview of a highly versatile forward and stable inverse algorithm for airborne, ground-based and borehole electromagnetic and electric data. exploration geophysics 46, 223–235, http://dx.doi.org/10.1071/eg13097 bockhorn, b., klint, k.e.s., jensen, m.b. & møller, i. 2015: use of geological mapping tools to improve the hydraulic performance of suds. water science and technology 71, 1492–1499, http://dx.doi. org/10.2166/wst.2015.125 brodic, b., malehmir, a., juhlin, c., dynesius, l., bastani, m. & palm, h. 2015: multicomponent broadband digital-based seismic landstreamer for near-surface applications. journal of applied geophysics 123, 227–241, http://dx.doi.org/10.1016/j.jappgeo.2015.10.009 christiansen, a.v., pedersen, j.b., auken, e., søe, n.e., holst, m.h. & kristiansen, s.m. 2016: improved geoarchaeological mapping with electromagnetic induction instruments from dedicated processing and inversion. remote sensing 8, 1022, http://dx.doi.org/10.3390/ rs8121022 christiansen, a.v. & auken, e. 2012: a global measure for depth of investigation. geophysics 77, wb171–177, http://dx.doi.org/10.1190/ geo2011-0393.1 doolittle, j.a. & brevik, e.c. 2014: the use of electromagnetic induction techniques in soils studies. geoderma 223–225, 33–45, http:// dx.doi.org/10.1016/j.geoderma.2014.01.027 jakobsen, p.r., hermansen, b. & tougaard, l. 2011: danmarks digitale jordartskort 1:25 000 version 3.1. danmarks og grønlands geologiske undersøgelse rapport 2011/40. copenhagen: geological survey of denmark and greenland. loke, m.h., chambers, j.e., rucker, d.f., kuras, o. & wilkinson, p.b. 2013: recent developments in the direct-current geoelectrical imaging method. journal of applied geophysics 95, 135–156, http:// dx.doi.org/10.1016/j.jappgeo.2013.02.017 neal, a. 2004: ground-penetrating radar and its use in sedimentology: principles, problems and progress. earth-science reiew 66, 261–330, http://dx.doi.org/10.1016/j.earscirev.2004.01.004 revil, a., karaoulis, m., johnson, t. & kemna, a. 2012: review: some low-frequency electrical methods for subsurface characterization and monitoring in hydrogeology. hydrogeology journal 20, 617–658, http://dx.doi.org/10.1007/s10040-011-0819-x viezzoli, a., christiansen, a.v., auken, e. & sørensen, k. 2008: quasi3d modeling of airborne tem data by spatially constrained inversion. geophysics 73, f105–113, http://dx.doi.org/10.1190/1.2895521 authors’ addresses k.e.s.k., geo, maglebjergvej 1, dk-2800 kgs. lyngby, denmark; e-mail kek@geo.dk. i.m., geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350 copenhagen k, denmark p.k.m & a.v.c., hydrogeophysics group, department of geoscience, aarhus university, c.f. møllers allé 4, dk-8000 aarhus c, denmark. geological survey of denmark and greenland bulletin 1, 75-114 75 the upper jurassic of europe: its subdivision and correlation arnold zeiss in the last 40 years, the stratigraphy of the upper jurassic of europe has received much attention and considerable revision; much of the impetus behind this endeavour has stemmed from the work of the international subcommission on jurassic stratigraphy. the upper jurassic series consists of three stages, the oxfordian, kimmeridgian and tithonian which are further subdivided into substages, zones and subzones, primarily on the basis of ammonites. regional variations between the mediterranean, submediterranean and subboreal provinces are discussed and correlation possibilities indicated. the durations of the oxfordian, kimmeridgian and tithonian stages are reported to have been 5.3, 3.4 and 6.5 ma, respectively. this review of the present status of upper jurassic stratigraphy aids identification of a number of problems of subdivision and definition of upper jurassic stages; in particular these include correlation of the base of the kimmeridgian and the top of the tithonian between submediterranean and subboreal europe. although still primarily based on ammonite stratigraphy, subdivision of the upper jurassic is increasingly being refined by the incorporation of other fossil groups; these include both megafossils, such as aptychi, belemnites, bivalves, gastropods, brachiopods, echinoderms, corals, sponges and vertebrates, and microfossils such as foraminifera, radiolaria, ciliata, ostracodes, dinoflagellates, calcareous nannofossils, charophyaceae, dasycladaceae, spores and pollen. important future developments will depend on the detailed integration of these disparate biostratigraphic data and their precise combination with the abundant new data from sequence stratigraphy, utilising the high degree of stratigraphic resolution offered by certain groups of fossils. this article also contains some notes on the recent results of magnetostratigraphy and sequence chronostratigraphy. keywords: europe, upper jurassic, oxfordian, kimmeridgian, tithonian, volgian, ammonite zonal and subzonal biostratigraphy and correlations, subdivision by non-ammonite fossil groups, chronometric data, magnetostratigraphy, sequence stratigraphy institut für paläontologie der universität erlangen-nürnberg, loewenichstr. 28, d-91054 erlangen, germany. present address: albert schweitzer strasse 19, d-91080 uttenreuth, germany. e-mail: arnold.zeiss@t-online.de geological survey of denmark and greenland bulletin 1, 75–114 (2003) © geus, 2003 76 contents subdivision and definition of stages: status and unsolved problems 78 the upper jurassic (malm) series 79 boundaries of the upper jurassic series 79 lower boundary (middle–upper jurassic series boundary) 79 upper boundary (jurassic–cretaceous system boundary) 80 upper jurassic stages – subdivision and correlation 81 oxfordian 81 lower boundary 82 substages 82 zones 84 correlation 85 chronometric data 85 kimmeridgian 85 lower boundary 86 additional remarks on lower kimmeridgian correlation 89 substages 89 zones 90 correlation 91 chronometric data 91 tithonian and volgian 91 lower boundary 92 substages 92 zones and subzones 93 correlation 95 chronometric data 96 biochronological importance of non-ammonite fossil groups: a review 96 invertebrate megafossil groups 96 cephalopods – other than ammonite conchs 96 aptychi 96 belemnites 96 bivalves 97 gastropods 97 brachiopods 97 echinoderms 97 corals (scleractinians) 97 sponges 97 vertebrate megafossils 97 invertebrate microfossils 98 foraminifera 98 radiolaria 98 ciliata 98 ostracodes 99 plant microfossils 99 dinoflagellata 99 calcareous nannofossils (coccoliths, nannolith groups) 99 charophyaceae 100 dasycladaceae 100 spores and pollen 100 magnetostratigraphy 100 sequence chronostratigraphy 100 acknowledgements 101 references 101 the term ‘upper jurassic’ (‘oberer jura’) was introduced by von buch (1839). arkell (1956) revived this name with only minor changes in its chronostratigraphic content. the term ‘upper jurassic’ in the sense of arkell (1956) was accepted by the first and second ‘colloque du jurassique’ at luxembourg in 1962 and 1967; only the stage name ‘purbeckian’ was eliminated, as it was considered to characterise merely a distinct lithofacies. this usage was followed by the five subsequent international symposia on jurassic stratigraphy at erlangen in 1984, lisbon in 1987, poitiers in 1991, mendoza in 1994 and vancouver in 1998. focus on the formal stratigraphic subdivision of the jurassic, and the upper jurassic in particular, is reflected in the series of key meetings since the early 1960’s (table 1). the term ‘malm’ was included in the recommendations of the first luxembourg colloquium in 1962 as an alternative term for the ‘upper jurassic’ (maubeuge 1964). although this term, like the term ‘tithonian’ (see below), is not based on a geographical site, it has been widely used since its introduction by oppel (1858, 1865). referring to the tithonian stage, arkell (1956, p. 8) wrote: “it is too late to abolish it after a hundred years of continuous use”; this also applies to the term ‘malm’. it is important to note that both ‘upper jurassic’ and ‘malm’ are chronostratigraphic terms; the latter, in particular, has frequently been used in a lithostratigraphic sense by some authors. at the first colloquium in luxembourg in 1962, a subdivision of the jurassic system into stages was proposed, the basic framework of which has survived to the present day. the stages were defined by their lower and upper ammonite zones. the recommendations of the first colloquium (maubeuge 1964) were thus a landmark in the history of international agreements concerning the subdivision of the jurassic system into series and stages. after a period of discussion following the publication of the resolutions of the luxembourg colloquia (maubeuge 1964, 1970), these proposals have been accepted 77 table 1. key events in upper jurassic stratigraphy since 1960 date place event reference 1962 luxembourg colloque du jurassique à luxembourg maubeuge 1964 1965 sofia vii congress, carpatho-balkan geological association cbga 1965 1967 luxembourg colloque du jurassique à luxembourg maubeuge 1970; brgm 1974 1967 moscow international symposium on upper jurassic stratigraphy anssr 1974 1969 london william smith symposium on jurassic geology 1969 budapest colloque du jurassique méditerranéen végh-neubrandt 1971 1973 neuchâtel colloque sur la limite jurassique–cretacé, lyon brgm 1975 1975 sofia international symposium on the jurassic–cretaceous boundary in bulgaria nikolov & sapunov 1977 1977 stuttgart international field meeting on the jurassic system of southern germany – this meeting zeiss 1977; ziegler 1977 initiated the reorganisation of the international subcommission on jurassic stratigraphy 1979 novosibirsk international colloquium on the upper jurassic and the jurassic–cretaceous boundary saks 1979 1984 erlangen international symposium on jurassic stratigraphy michelsen & zeiss 1984 1984 sümeg meeting of the working group for the jurassic–cretaceous boundary fülöp 1986 1987 lisbon 2nd international symposium on jurassic stratigraphy rocha & soares 1988 1987 international field meeting on jurassic–cretaceous boundary problems at menner 1990 the northern caucasus 1988 zaragoza 1st oxfordian working group meeting, zaragoza – iberian chain meléndez 1990 1990 basel 2nd oxfordian working group meeting, basel and jura range of northern switzerland gygi 1990b 1991 poitiers 3rd international symposium of jurassic stratigraphy cariou & hantzpergue 1994 1992 warsaw joint meeting of the oxfordian and kimmeridgian working groups atrops et al. 1993a 1993 london w. j. arkell symposium of jurassic geology taylor 1996 1994 mendoza 4th international symposium on jurassic stratigraphy riccardi 1996 1994 lyon 4th oxfordian and kimmeridgian working groups meeting, lyon and atrops & meléndez 1994a south-eastern france basin 1997 warsaw oxfordian (jurassic) meeting in poland glowniak et al. 1997 1998 vancouver 5th international symposium on the jurassic system pálfy 1998, hall & smith 2000 worldwide, the only exception being that in the former soviet union the callovian has been considered to belong to the upper jurassic (see krymholts et al. 1988), while in the rest of the world the callovian is included in the middle jurassic. however, following a decision by the interdepartmental stratigraphic commitee in 1989, the callovian is also now considered in russia to belong to the middle jurassic (zhamojda 1991). according to the most recent estimates, the late jurassic had a duration of a little more than 15 million years according to gradstein et al. (1994, 1995; ogg 1995), or 12 million years (+5.6/-7.3) according to pálfy et al. (1998). the data of gradstein et al. (1995) have been used in figures 2, 4 and 5 of this paper; on this basis each of the three upper jurassic stages has an average length of 5 million years, while zones and subzones have approximate durations of 700 000 and 300 000 years, respectively. each subzone comprises at least three horizons, each of which has an approximate duration of 100 000 years. subdivision and definition of stages: status and unsolved problems on the basis of the recommendations of the first luxembourg colloquium (maubeuge 1964), the upper jurassic series was subdivided into four stages for the boreal and subboreal regions: oxfordian, kimmeridgian (sensu anglico), portlandian (sensu anglico) and volgian, and three for the submediterranean and mediterranean regions: oxfordian, kimmeridgian (sensu gallico equivalent to ‘crussolian’) and tithonian (equivalent to ‘danubian’ and ‘ardescian’). in the following decades, there has been much confusion as the kimmeridgian and portlandian stages have often been used differently in different parts of europe. in 1990, a formal vote of the international subcommission on jurassic stratigraphy (isjs) led to the decision to use stages with approximately the same vertical age ranges and uniform names in both regions: kimmeridgian (sensu gallico) and tithonian (zeiss 1991a). with regard to the still unresolved correlation problems between the boreal and mediterranean provinces, it was agreed that the volgian can be used as an alternative stage for the tithonian in subboreal and boreal regions (fig. 1). the main problem which remained to be solved was the definition of the lower boundary of each stage. to date, no stage has a type locality and a defined lower boundary (global stratotype section and point, or gssp) formally accepted by the international commission on stratigraphy (ics). there are of course a lot of proposals, but they have not been validated according to the guidelines and rules of the ics (cowie et al. 1986; remane et al. 1996). the most intractable problems are to find isochronous levels in submediterranean (and mediterranean) and subboreal (and boreal) europe for the lower boundary of the kimmeridgian stage and for the upper boundary of the tithonian (volgian) stage. as the former is of particular significance for upper jurassic subdivision and correlation, it will be treated here in some detail (see below). other problems are the unification of the differing subdivisions of stages into substages, correlation of the zones of each stage between the different areas of europe and the development of better correlation charts from the boreal regions to the mediterranean areas. provisional correlation charts on a zonal and subzonal level for each stage of the european upper jurassic are presented here (see figs 2–5). zones and subzones are used here as chronozones following the international stratigraphic guide (salvador 1994); originally, many of them were defined as biozones whereas others were used as standard zones, standard chronozones or biochronological standard zones, i.e. only the base is defined while the top is defined by the base of the next overlying unit (callomon 1965, 1984a, 1994). problems arise, however, due to inconsistent usage of the term ‘standard zone’. in northwest europe, standard zones are mostly used following the concept of callomon (1994), whereas in central and southern europe, standard zones are often synonymous with biostratigraphic zones for use in biochronology (cariou & hantzperque 1997). it is often difficult, therefore, to determine in which meaning ‘standard zones’ are used. the problems encountered in moving from biostratigraphic field data to biochronological interpretations have been discussed recently by remane (1991). the term biochronological zone is now used by many authors instead of chronostratigraphic zone, if the zone is based on fossil data. as the ultimate subdivision of biochronology, french authors use the term ‘biohorizon’ (e.g. enay 1997); their concept is therefore ‘sensiblement different’ from the pure biostratigraphic horizon concept of j.h. callomon (dommergues 1997). no attempt has been made here to correlate ammonite faunal horizons due to the variable nature of the published research on the upper jurassic in the various sedimentary basins of europe. where necessary, however, correlations of horizons are discussed in the text. 78 for an example of such horizon correlations, the reader is referred to the work of callomon (1984c) on the upper jurassic of north america (this study also covers the european amoeboceras subdivision). although attempts have been made to generalise horizons for the whole ‘domaine tethysien’ and ‘domaine boréal’ (cariou et al. 1997; hantzperque et al. 1997), these appear premature and of little practical use, given the present state of knowledge. in europe, subdivisions down to the level of ammonite faunal horizons have been proposed for several sedimentary basins. such studies include that of the upper jurassic of east greenland by callomon & birkelund (1980, 1982; birkelund & callomon 1985), the lower kimmeridgian of southern england by birkelund et al. (1983), the kimmeridgian of spitsbergen by wierzbowski (1989), the kimmeridgian of the barents sea by wierzbowski & smelror (1993), the lowermost oxfordian of northern france by vidier et al. (1993), the upper oxfordian, kimmeridgian and lower tithonian of western france by hantzperque (1989), the lowermost oxfordian of south-east france by fortwengler & marchand (1994b), the lower kimmeridgian of south-east france by atrops (1982), the upper oxfordian, upper kimmeridgian and lower tithonian of south-west germany by schweigert (1994, 1995a, b, 1996a, b; schweigert & callomon 1997), the oxfordian– kimmeridgian of poland (matyja & wierzbowski 1997), the middle volgian of central poland (kutek 1994) and the oxfordian of north-east spain by cariou et al. (1991a) and meléndez & fontana (1993). the ‘faunal horizon’ approach clearly represents a method for increasing precision in correlation and dating in the future, when the data from the various sedimentary basins reach the necessary standard. it is already proving useful in deciphering the history of basin deposits at a resolution that was hitherto impossible; such data, in particular, allow us to date events more precisely and to determine the ‘completeness’ of the sedimentary record i.e. to identify accurately the position and duration of hiatuses. a prerequisite is, however, that it is possible to reconstruct the complete succession of faunal horizons by correlating individual local successions. in this context, it is worth mentioning that the methods of jurassic stratigraphy and high-resolution geochronology have been discussed in detail by callomon (1984a, b, 1994, 1995), page (1995), corna et al. (1997), blau (1998) and blau & meister (2000); formal aspects were covered by remane (1996). the upper jurassic (malm) series (fig. 1) in this paper, subdivision and correlation of the upper jurassic series have been carried out mainly using ammonites. other fossil groups are reviewed briefly, however, with request to their biochronologic resolution and correlation potential. many papers have been published on upper jurassic ammonites and their chronostratigraphic resolution (see detailed discussion below). for a broad overview, the reader is referred to the papers of cariou et al. (1997), geyssant (1997) and hantzpergue et al. (1997) for western europe and the mediterranean. other important, partly regional compilations and revisions have been published by sapunov (1979), donovan et al. (1981), krymholts et al. (1988), malinowska et al. (1988), enay et al. (1994) and schlegelmilch (1994). boundaries of the upper jurassic series lower boundary (middle–upper jurassic series boundary) the lower boundary of the oxfordian stage is rather well-defined by ammonite zones and subzones and only requires more precise definition with respect to the lowermost faunal horizon, which then would characterise the beginning of the lowermost subzone (and zone) of the stage. furthermore, it appears that the lower boundary is approximately (on a subzonal level) the same in boreal and mediterranean areas. once the type faunal horizon has been chosen, then the problem of the type locality for the boundary will also have been solved. at present, this boundary lies in france between the uppermost horizon of the quenstedtoceras lamberti zone of the upper callovian substage (the cardioceras paucicostatum horizon) and the lowermost horizon of the q. mariae zone; this was first named in france after peltoceratoides elisabethae (fortwengler & marchand 1994a), but afterwards was changed to hecticoceras (brightia) thuouxense (fortwengler & marchand 1994b, c), a species described only recently (fortwengler et al. 1997). in dorset, however, cardioceras cf. woodhamense and c. woodhamense are found in the lowermost levels of the q. mariae zone (callomon & cope 1996), whereas in north-west france, c. woodhamense has been collected only in the third horizon of the q. mariae zone (vidier et al. 1993). in south-east france, this horizon is only recognised tentatively. these different faunal horizons all lie 79 in the cardioceras scarburgense subzone, the lower subzone of the q. mariae zone, so that the age difference of these horizons (if any) should not be too large. a vote by the callovian/oxfordian boundary working group in 1995 resulted in a preference for a type locality in south-east france, with the consequence that the oxfordian would begin with the h.(b.) thuouxense horizon (see above), but a final decision was not taken (meléndez 1995; meléndez et al. 1998). upper boundary (jurassic–cretaceous system boundary) in accordance with the decision of the isjs (see above), there are two alternative stages for the uppermost part of the jurassic system: tithonian and volgian. as they differ in duration, the boundary may be drawn at two different levels, i.e. there are two variants of the jurassic–cretaceous boundary. accordingly, the members of the former jurassic–cretaceous boundary working group agreed to work provisionally with two boundaries (remane 1986; remane et al. 1986; zeiss 1986). 1. in mediterranean and submediterranean europe, the boundary is placed between the top of the tithonian stage (top of durangites vulgaris zone and/or of calpionellid zone a) and the base of the berriasian stage (base of berriasella jacobi zone s.l. (= berriasella jacobi and pseudosubplanites grandis subzones or pseudosubplanites euxinus zone) and/or base of calpionellid zone b). 2. in subboreal and boreal europe, the boundary lies between the upper volgian (top of craspedites nodiger or chetaites chetae zone) and the ryazanian or ‘boreal berriasian’ (base of chetaites sibericus, rjasanites rjasanensis or runctonia runctoni zone) (rawson et al. 1978; kejsi et al. 1988; sey & kalacheva 1993a). 80 upper upper upper upper upper upperupper upper middle middle middle middle middle middle middle lower lower lower lower lower lower lowerlowerlower abnormis scythicus klimovi autissiodorensis mutabilis baylei glosense plicatilis mariae semiforme hybonotum beckeri acanthicum platynota bimammatum bifurcatus transversarium mariae tithonian kimmeridgian oxfordian tithonian (volgian) stagesstages substagessubstages germany france submediterranean province subboreal province basal zones of substages basal zones of substages kimmeridgian oxfordian fig. 1. subdivision of the upper jurassic series of europe into stages, substages and zones. substage usage varies in the literature, dependent on author; those indicated are only examples. in the first case, the type locality should be best selected in south-eastern france, where the ardescian substage (upper tithonian) and the berriasian stage were originally described. subsequent studies have revealed that the sequences are not complete at the base, however, so that it has been suggested that the best sections illustrating the jurassic–cretaceous boundary beds and their fauna are situated in southern spain (enay & geyssant 1975; tavera 1985; tavera et al. 1994; enay et al. 1998a, b). in the second case, the boundary should correspond to the base of the berriasella boissieri zone in the mediterranean area. thus, the upper volgian substage corresponds to the lower berriasian (zeiss 1974, 1979, 1983, 1986; rawson et al. 1978; hoedemaker 1990; sey & kalacheva 1993a; w.a. wimbledon in: callomon & cope 1996), and is not equivalent to the upper tithonian as mesezhnikov (1988) and other authors have assumed. in a recent review of the berriasian stage, hoedemaker (1994) stated that the jurassic–cretaceous boundary is typically placed at one of two different levels, either at the base or at the top of the jacobi chronozone: “investigators of jurassic stratigraphy prefer the lower of these two boundaries, investigators of the cretaceous stratigraphy the upper” (hoedemaker 1994, p. 12). at the same time, there has also been an attempt to trace the jurassic–cretaceous boundary based on geomagnetic anomalies from the tethys to southern england (ogg et al. 1994). in the tethyan–atlantic faunal realm, the top of magnetic polarity reversal m19r approximately coincides with the tithonian–berriasian boundary in the mediterranean area. this reversal is difficult to place precisely in england, but it seems to be situated in the lowermost purbeck beds. if so, it would demonstrate once again that the ‘upper volgian’ (casey 1973) or ‘upper portlandian’ of england (wimbledon 1980), i.e. the zonal sequence subcraspedites primitivus – subcraspedites lamplughi, overlaps with the lower berriasian. wimbledon (1980) also included the ‘upper volgian’ zones of casey (1973) in the ‘portlandian’ of britain, thus extending the stage upwards by three further zones (termed here ‘upper portlandian’). in a recent compilation chart, w.a. wimbledon (in: callomon & cope 1996) correlated these ‘upper portlandian’ zones and the upper volgian zones of the russian platform with parts of the lower berriasian. in poland, the jurassic–cretaceous boundary has been traced by joint studies of ostracodes and ammonites (marek et al. 1989) whereby the upper tithonian and the lower part of the lower berriasian could be recognised as well as the upper berriasian (= ‘ryazanian’). in a recent paper (marek & shulgina 1996), the ammonites of the berriasian (ryazanian) were considered to belong to the interval upper occitanica – lower boissieri zones. in a recent development, the interdepartmental stratigraphic committee of russia (isc) approved the following resolutions of its commissions on the jurassic and cretaceous systems (rostovtsev & prozorovskiy 1997, p. 48). “1. to draw the jurassic–cretaceous boundary in the boreal realm between the middle and upper substages of the volgian, and not …… as …… earlier adopted in russia (1978). this boundary mainly corresponds to the tithonian/berriasian boundary in tethyan realm (colloque lyon–neuchâtel, 1975). correspondingly, the lower volgian in the whole correlated with the lower and middle tithonian; the middle volgian, with the upper tithonian; the upper volgian, with two lower zones of the berriasian (jacobi/grandis and occitanica). 2. to transfer the volgian stage in its former range to the category of regional stratigraphic units (regional stage). to distinguish as chronostratigraphic units in the boundary part of the jurassic and cretaceous scale of russia only tithonian and berriasian.” these resolutions, which were precipitated by the work of sey & kalacheva (1993a), confirmed the earlier opinions of many authors concerning upper jurassic/lower cretaceous correlations. it is clear that general consensus has not yet been reached; it is assumed, however, that the present tithonian–berriasian boundary is not suitable for global correlation. it may be preferable, therefore, to return to an old proposal: to define the jurassic–cretaceous boundary at the base of the b. boissieri zone, where many guide fossils of different groups are available for correlation. recent studies in the caucasus area by remane (1997) are supportive of this proposal. upper jurassic stages – subdivision and correlation oxfordian (figs 2, 3) the colloquium at luxembourg in 1962 (maubeuge 1964, p. 85) came to the resolution “...... that it seemed necessary to return to the original sense of this stage 81 [the oxfordian] as defined by a. d’orbigny and given precision by w.j. arkell (1956)”. the ‘base’ was indicated to be the ‘zone of quenstedtoceras mariae’ and the ‘top’ the ‘zone of ringsteadia pseudocordata (= zone of idoceras planula), (= zone of epipeltoceras bimammatum)’. it was recommended that other stage and substage names then still in use, e.g. the argovian (marcou 1848), rauracien (greppin 1867), sequanian (marcou 1848) and the lusitanien (choffat 1885; haug 1910) should be abandoned. these stages had been interpreted in different ways so that continued usage would have created only more confusion. subsequent studies (e.g. enay 1980a; gygi & persoz 1986; enay et al. 1988) demonstrated the validity of this resolution. lower boundary see discussion above. substages although the oxfordian has been subdivided into three substages, lower, middle and upper oxfordian, full agreement has not been reached on the zonal content of these substages and the position of their boundaries (callomon 1988, 1990, fig. 10; meléndez & fontana 1993, fig. 5; wright 1996a). the subdivision is thus essentially informal but the substages are capitalised in 82 switzerlandfrance, spain s. germany england greenland, scotland, svalbard galar (gigantoplex, grandiplex) galar planula galar planula (grandiplex) planula proteron pl an ul a pl an ul a luciaeformis (wartae) parandieri parandieri nunningtonense cautisnigrae serratum serratum koldewayense regulare rosenkrantzi variocostatum caledonica pseudoyo parandieri schilli schilli schilli rotoides stenocycloides stenocycloides grossouvrei grossouvrei antecedens antecedens antecedens tenuiserratum glosense (alternoides) glosense ilovaiskyi tenuiserratum blakei pu m ilu s c au tis ni gr ae antecedens vertebrale vertebrale vertebrale maltonense vertebrale cordatum cordatum cordatum cordatum cordatum costicardia costicardia costicardia costicardia costicardia bukowskii bukowskii bukowskii bukowskii bukowskii praecordatum praecordatum praecordatum praecordatum praecordatum scarburgense scarburgense scarburgense scarburgense scarburgense plicatilis pl ic at il. pl ic at ili s c or da tu m c or da tu m c or da tu m cordatumcordatum m ar ia e m ar ia e m ar ia e mariaemariae a lte rn an s o va le ba uh in i r in gs te ad ia densiplicatum densiplicatum transversarium transversarium tr an sv er sa ri um tr an sv er sa ri um bifurcatus bifurcatus hypsel. hauffianum bimammatum bimammatum bimammatumbimammatum bi m am m at um ps eu do co rd at a ba yl ei r av ni (s ub -) bo re al o x. /k i. bo un da ry pseudocordata evoluta densicostata baylei ? ? bauhini (? bayi) bauhini tonnerense praecursor praecursor berrense berrense hauffian. semimammatum bi fu rc at us submediterranean subboreal boreal 15 9. 4 (± 3 .6 ) 15 4. 1 (± 3 .2 ) u pp er m id dl e lo w er o xf or di an fig. 2. a tentative correlation chart for the oxfordian stage in europe (thick lines indicate periods during which correlation is difficult). modified after zeiss (1984), mesezhnikov (1988), cariou et al. (1991b), wright (1996a, b), matyja & wierzbowski (1997, 1998), schweigert & callomon (1997) and gygi (2000a, b). this paper, following common usage. an example of the ongoing debate is the inclusion of the d. bifurcatus zone in the middle or upper oxfordian; this zone was introduced by enay (1966) as the upper subzone of the g. transversarium zone but was later considered as the lowermost zone of the upper oxfordian (cariou et al. 1971). preference is given here to a subdivision in which the d. bifurcatus zone is included in the middle oxfordian (fig. 1) as has also been proposed by meléndez (1989), cariou & meléndez (1990), cariou et al. (1991a) and gygi (2000a) although not followed by cariou et al. (1991b, 1997). while the lower and middle substages have the same lower boundaries in submediterranean and subboreal europe, the position of the lower boundary of the upper substage differs. in boreal europe, it has been drawn at three different levels (wright 1996a, fig. 6). the solution to draw it at the base of the a. glosense zone is well-known (sykes & callomon 1979; wright 1980); it would correspond to the base of the p. luciaeformis subzone in the g. transversarium zone, i.e. the boundary would be drawn around one and a half zones deeper than in the submediterranean subdivision. it seems preferable to draw the boundary at the lower boundary of the a. rosenkrantzi zone, corresponding approximately to the lower boundary of the upper oxfordian both in submediterranean europe (base of e. bimammatum zone) and in subboreal europe (base of r. pseudocordata zone), although the latter lies somewhat deeper (matyja & wierzbowski 1997, fig. 4). additional literature references pertinent to the subdivision of the oxfordian stage are enay (1963, 1966), zeiss (1966), sequeiros (1974), sapunov (1976), gygi (1977, 1986, 1990a, 2000a, b, c), wierzbowski (1978), enay & meléndez (1984), a. zeiss (in: enay & meléndez 1984), cariou & meléndez (1990), malinowska (1991), meléndez & fontana (1993), schweigert (1995a, b), fözy & meléndez (1996), matyja & wierzbowski (1997, 1998), groiss et al. (2000) and schweigert & callomon (1997) for the submediterranean and mediterranean provinces, and sykes & callomon (1979), wright (1980, 1996a, b) and mesezhnikov (1988) for the subboreal and boreal provinces. mönning & bertling (1995), mönning (1998) and gramann et al. (1997) have presented interesting and useful reviews of the ammonite succession in northern germany. 83 submediterranean standard (cariou et al. 1991b) spain – north africa (sequeiros 1974; cariou et al. 1991b) poland (tarkowski 1990; matyja & glowniak 1994) bulgaria (sapunov 1976) transversarium plicatilis riazi antecedens antecedens promiscuus oculatum helenaepaturattensis baccatum spixi wartae antecedens episcopalis renggeri athletoides riazi antecedens cordatum costicardia bukowskii praecordatum scarburgense vertebrale (tenuicostatum) rotoides schilli luciaeformis parandieri paturattensis claromontanus claromontanus paturattensis paturattensis plicatilis minax magnouatius mazuricuscordatum mariae transversarium fig. 3. a tentative correlation chart for some alternative subdivisions of parts of the oxfordian stage in mediterranean and submediterranean europe. zones the zonal and subzonal subdivision of the lower oxfordian substage was established by arkell (1941) using quenstedtoceras mariae and cardioceras cordatum as index species; it can be used over large areas of northern and central europe (fig. 1) and is also applicable in the dauphinois basin of south-east france as recently demonstrated by fortwengler & marchand (1994a; fortwengler et al. 1995). in southern europe, a variety of subdivisions exist; at least three distinct subdivisions testify to the difficulties in erecting a generally accepted zonal scheme if cardioceratids are missing. in such cases, peltoceratids (peltomorphites, peltoceratoides and parawedekindia), oppeliids (taramelliceras, popanites and creniceras) and perisphinctids (otosphinctes, perisphinctes, prososphinctes and properisphinctes) are important guide fossils (see fig. 3), e.g. taramelliceras minax, t. spixi, t. baccatum, t. oculatum, popanites paturattensis in poland (tarkowski 1990), peltomorphites athletoides and creniceras renggeri in bulgaria (sapunov 1976), and prososphinctes mazuricus and p. claromontanus in spain (aurell et al. 1990). from the base of the middle oxfordian, perisphinctids and peltoceratids become the dominant ammonite groups with respect to index fossils at the substage level in the submediterranean and subboreal provinces. the first aulacostephanids (decipia) also appear at this level. the perisphinctes plicatilis, gregoryceras transversarium and dichotomoceras bifurcatum zones make up the middle oxfordian substage in the submediterranean area, the perisphinctes plicatilis, p. pumilum and p. cautisnigrae zones are representative of the subboreal province. boreal indexes are cardioceras tenuicostatum and c. tenuiserratum, amoeboceras glosense and a. serratum. the correlation between subboreal perisphinctid and amoeboceratid zones was well demonstrated by wright (1996b). there is a difference in the usage of the p. plicatilis and g. transversarium zones in submediterranean europe. although gygi & marchand (1982) replaced the basal c. vertebrale subzone with the c. densiplicatum zone and included the p. antecedens subzone in the g. transversarium zone, subsequent authors have not followed the arguments of these authors and have continued to use the p. plicatilis zone in the sense of cariou et al. (1991a, b), i.e. with c. vertebrale and p. antecedens subzones (e.g. meléndez & fontana 1993, fig. 4; cariou et al. 1997). cariou et al. (1991a) defined the g. transversarium zone to contain the p. parandieri, p. luciaeformis, l. schilli and p. rotoides subzones. in a more recent publication, gygi (1995) included in the lower part of the g. transversarium zone not only the p. antecedens subzone but also the c. densiplicatum subzone i.e. the whole p. plicatilis zone (following the original usage of oppel & waagen 1866; r.a. gygi, personal communication 1997). in further contributions to the upper jurassic of switzerland (gygi 2000b, c), the g. transversarium zone is subdivided into the c. densiplicatum, p. antecedens, and p. luciaeformis subzones; the overlying d. bifurcatus zone contains in its lower part the l. schilli subzone, which is considered in spain and france to represent the upper part of the g. transversarium zone (see above). the main reason for these differences is the occurrence of l. schilli in switzerland above the vertical range of g. transversarium. the upper oxfordian substage in submediterranean europe comprises the epipeltoceras bimammatum, idoceras planula and sutneria galar zones. considering the new correlations of wright (1996a), matyja & wierzbowski (1997) and cariou et al. (1997), ringsteadia pseudocordata would be the corresponding index fossil for the subboreal province, whereas amoeboceras rosenkrantzi would be the index fossil for boreal europe. the amoeboceras serratum zone of malinowska (1991) contains epipeltoceras (uhligi group) and ringsteadia salfeldi thus indicating, at least partly, equivalence with the lower e. bimammatum zone (e. hypselum subzone); this demonstrates that the a. serratum zone of this author is younger in age than the a. serratum zone of sykes & callomon (1979). the a. regulare subzone of malinowska (1991) seems to represent the upper e. bimammatum and perhaps the lowermost i. planula zones, while the a. lineatum subzone apparently corresponds to the rest of the i. planula zone and the s. galar zone. the most difficult problems associated with these upper oxfordian zones concern their correlation in the subboreal and submediterranean schemes; this aspect is discussed in detail below. some minor problems may be caused by the different hierarchical status of zones and subzones in the subboreal and boreal provinces. for example, atrops et al. (1993b) recognised the a. regulare, a. rosenkrantzi and a. bauhini zones, whereas malinowska (1991) subdivided the r. pseudocordata zone into the a. regulare and a. lineatum subzones or, in boreal europe, into the a. regulare and a. rosenkrantzi subzones. however, comparing the correlation chart of malinowska (1991, table 3) with that of matyja & wierzbowski (1997, fig. 3), it becomes evi84 dent that the p. pseudocordata zone of malinowska corresponds only to the upper part of the a. regulare, the a. rosenkrantzi and the p. baylei zones. another example is the variable status of a. bauhini as an index species. there is the a. bauhini horizon in the upper e. bimammatum zone equivalent to the p. densicostata horizon (schweigert & callomon 1997), the a. bauhini subzone of the a. rosenkrantzi zone (sykes & callomon 1979; cariou et al. 1997), equivalent to the p. baylei zone of birkelund & callomon (1985), and the a. bauhini zone. although initially equivalent to the p. densicostata horizon (wierzbowski & smelror 1993), the a. bauhini zone was expanded by matyja & wierzbowski (1997, 1998) to correlate with the uppermost p. pseudocordata zone and nearly the whole p. baylei zone on the one hand and with the whole i. planula zone and uppermost e. bimammatum zone on the other; a little more restricted was the a. bauhini zone of schweigert & callomon (1997), who excluded the s. galar subzone of the i. planula zone (see below). correlation there have been many proposals and attempts to correlate the zonal subdivisions of the oxfordian of mediterranean, submediterranean, subboreal and boreal areas of europe; the most important ones have been already discussed in the text above (see figs 2, 3). further informative compilations have been presented by enay & meléndez (1984), mesezhnikov (1988), cariou et al. (1991a, b; 1997), malinowska (1991), aleynikov & meledina (1993), meléndez & fontana (1993), schweigert (1995b), wright (1996a, b) and matyja & wierzbowski (1997). in short, correlation within the lower oxfordian is possible over wide regions of boreal, subboreal and submediterranean europe, but becomes difficult on approaching the mediterranean area. at the base of the middle oxfordian, ammonites of the perisphinctes plicatilis zone provide the last possibility for long-distance correlation. higher up in the middle oxfordian, zonal correlations become more and more difficult, best illustrated by the charts of j.h. callomon (in: wright 1980), enay & meléndez (1984) and cariou et al. (1991b; see also fig. 2). the divergent views are also well-documented by the tables of malinowska (1991), wright (1996a), cariou et al. (1997) and matyja & wierzbowski (1997). the problems of upper oxfordian correlation, concentrated mainly on the correspondence of the e. bimammatum, i. planula and s. galar zones to the r. pseudocordata, p. baylei, a. regulare, a. rosenkrantzi and a. bauhini zones, are under discussion (wierzbowski 1991; atrops et al. 1993b; atrops & meléndez 1994b; schweigert 1995a, b; cariou et al. 1997; matyja & wierzbowski 1997; schweigert & callomon 1997). this aspect is especially relevant to the oxfordian– kimmeridgian boundary problem and is therefore discussed in more detail below. chronometric data the duration of the oxfordian stage is estimated at 5.3 ma (gradstein et al. 1995; ogg 1995; ogg & gutowski 1996); for precise data, see figure 2. kimmeridgian (fig. 4) following the luxembourg recommendations of 1962 and 1967 (maubeuge 1964, 1970), two possibilities existed with respect to usage of the kimmeridgian stage, namely either a long version (‘sensu anglico’) or a short version (‘sensu gallico’), both with differing zonal content and boundaries (see below). use of two different versions of the kimmeridgian evoked much confusion in following years and led to endless discussion. therefore a vote of the international subcommission on jurassic stratigraphy (isjs) on this question was arranged in 1990, simultaneously with the vote on the tithonian stage (see below); the members of the isjs voted for a ‘short’ version of the kimmeridgian stage (i.e. ‘sensu gallico’). this meant that in future the upper boundary of the kimmeridgian stage should be coincident with the lower boundary of the tithonian stage and its boreal equivalent, the volgian (zeiss 1991a). the lower boundary of the stage, however, remained ambiguous (see below). because of the still unresolved problems at the oxfordian–kimmeridgian boundary, the lower boundary of the kimmeridgian stage is drawn in this paper at the base of the sutneria platynota zone, following the above-mentioned adoption of a short kimmeridgian stage (i.e. ‘sensu gallico’ or according to the ‘continental’ concept; enay 1980b). the working group of the oxfordian–kimmeridgian boundary is mandated to finally define the boundary at a level which allows farreaching correlations and corresponds to the resolutions of the international commission on stratigraphy (ics); see also the discussions by wierzbowski (1999, 2001). 85 until such a definition has been taken by the oxfordian–kimmeridgian boundary working group, voted on by the isjs and approved by isc, it seems useful to maintain the traditional boundaries in both biogeographic provinces, and it is premature to draw the oxfordian–kimmeridgian boundary in the submediterranean area in the upper part of the e. bimammatum zone (cf. gygi 2000a, b). lower boundary as the luxembourg recommendations made it possible to select between two distinct versions of the kimmeridgian stage, the lower boundary was also defined twofold. in subboreal regions of europe, the boundary was drawn at the lower boundary of the pictonia baylei zone, whereas in submediterranean regions it was placed at the base of the sutneria platynota zone (maubeuge 1964, p. 85–86). at that time, it was supposed that both boundaries were more or less isochronous (ziegler 1964), although doubts remained (e.g. zeiss 1965; cariou et al. 1971). with the publication of sykes & callomon (1979), new impetus was given to further studies, which have suggested that the assumed time equivalence is erroneous or, at best, only partially true (matyja & wierzbowski 1988; wierzbowski 1991; atrops et al. 1993b; schweigert 1995a, b). the main reasons for this view were the discovery of new amoeboceras faunas by these authors 86 beckeri/ pressulum beckeri eudoxus eudoxus mutabilis cymodoce cymodoce achilles chatelaillonensis cymodoce baylei bayi subkitchini modestum ruepellense mutabilis lallierianum orthocera caletanum contejeani autissiodorensis irius mutabilis autissiodorensis (volgensis) autissiodorensis (taimyrense) elegans kochi k oc hi n or ve gi cu m autissiodorensis eudoxus eudoxuscavouri acanthicum acanthicum divisum hypselocyclum platynota polygyratus desmoides guilheradense hippolytense lothari crusoliense uhlandi balderum linealis attenuatus eulepidus liparum/ schilleri eudoxus caletanum subeumela setatum ulmense submediterranean s. germany mediterranean n. italy (s. alps) biome franco-germanique w. france subboreal great britain boreal n. europe mixed poland herbichi strombecki silenum trenerites raschi stenonis divisum uhlandi fallax subborealis a ut is si od or en si s m ut ab ili s c ym od oc e k itc hi ni acanthicum divisum hypselocyclum platynota u pp er m id dl e k im m er id gi an lo w er 15 0. 7 (± 3. 0) 15 4. 1 (± 3. 2) fig. 4. a tentative correlation chart for the kimmeridgian stage in europe (thick lines as in fig. 2). modified after zeiss (1965), atrops (1982), sarti (1988), hantzpergue et al. (1991), wierzbowski & smelror (1993), kutek & zeiss (1994), schweigert & zeiss (1994) and matyja & wierzbowski (1997, 1998). and a re-evaluation of salfelds (1915) cardioceras paper as well as that of koerner (1963), particularly with respect to their remarks concerning the type locality and possible type horizon of cardioceras (= amoeboceras) bauhini. the discussion of wierzbowski (1991) concerning the range of the genus ringsteadia in poland is also important in this context. it soon became evident that amoeboceras bauhini has its type horizon just below the upper boundary of the e. bimammatum zone, (see a. zeiss in: enay & meléndez 1984). the studies of schweigert (1995a, b; schweigert & callomon 1997) resulted in similar conclusions, but led to a more precise faunal horizon subdivision of the upper oxfordian in württemberg, sw germany and to better correlation possibilities with england, with respect to the a. bauhini and the a. bayi (?= a. subtilicaelatum) horizon. some problems remained unsolved, however: 1. does the a. bauhini horizon of southern germany represent the same time interval as the beds bearing a. bauhini in england, scotland and the barents sea? or is there a difference, and the vertical range of this species is different in these two areas? what is the situation in poland, representing an intermediate region? 2. does the a. subtilicaelatum horizon of southern germany represent the same time interval as the a. bayi horizon in england? or is there also a difference in the vertical range of these species in different parts of europe? 3. which units in the subboreal realm correspond to the succession from the base of the i. planula zone (with three or four faunal horizons) and the top of the lower s. galar zone, which in submediterranean europe occurs between the a. bauhini and the a. subtilicaelatum (?= a. bayi) horizon? it is not easy to answer these questions given the present state of knowledge; the following points are pertinent prior to discussion of these problem areas. the usage of a. bauhini as an index ammonite began with its introduction by sykes & callomon (1979) as a subzone of the a. rosenkrantzi zone (uppermost oxfordian); its stratigraphic position was subsequently revised by birkelund & callomon (1985), who regarded the a. bauhini subzone and the p. baylei zone (lower kimmeridgian) as approximate equivalents. one year prior to this latter publication, a. zeiss (in: enay & meléndez 1984, fig. 6) had used a. bauhini informally as a zonal index in a correlation chart to show its approximate correspondence with the i. planula zone sensu lato; this view was also held by atrops et al. (1993b) and matyja & wierzbowski (1997, 1998). wierzbowski & smelror (1993) established the a. bauhini zone formally and suggested that it was equivalent to only the lower part of the p. baylei zone (the p. densicostata horizon); in more recent papers, matyja & wierzbowski (1994, 1995, 1997, 1998) provided charts showing the correlation between the a. bauhini zone and the i. planula zone sensu lato as well as with the p. baylei zone (with the exception of the uppermost part). finally, in southern germany, an a. bauhini horizon was described by schweigert (1995b; schweigert & callomon 1997) in the upper part of the t. hauffianum subzone (uppermost e. bimammatum zone); the latter authors correlated the boreal a. bauhini zone with the i. planula zone sensu stricto, whereas the s. galar zone was correlated with the amoeboceras kitchini zone. the amoeboceras bayi horizon was introduced by birkelund & callomon (1985) in the upper part of the p. baylei zone, whereas wierzbowski & smelror (1993) reported the species at the base of their a. subkitchini subzone. atrops et al. (1993b) found the species, or closely related forms, in the sutneria platynota zone of the submediterranean area. schweigert (1995b) established an a. subtilicaelatum horizon in the uppermost part of the sutneria galar zone, assuming that a. bayi is only a variant of a. subtilicaelatum, which would then have priority. this conflicts with the opinion of salfeld (1915), that a. lineatum and a. subtilicaelatum are very close and perhaps synonymous. schweigert (1995b) also assumed that many specimens determined earlier as ‘a. bauhini’ belong in reality to a. bayi. to verify these assumptions, a comprehensive re-evaluation of the upper oxfordian – lower kimmeridgian amoeboceras species complex (a. bauhini – a. bayi – a. subtilicaelatum – a. lineatum) would be necessary. such a study should also illustrate the variation within each species in time and space (see, for example, klieber 1981; birkelund & callomon 1985; matyja & wierzbowski 1988, 1994; schweigert & callomon 1997). for the time interval of the i. planula zone, malinowska (1991) established the a. lineatum subzone in poland. it was introduced as the upper subzone of the r. pseudocordata zone, but the precise correlation with other areas is not clear; from the list of fossils one would conclude that the s. galar zone is not present. however, as a sutneria sp. (of the galar/praecursor group?) is mentioned in the text but not figured, a deci87 sion is difficult; its low stratigraphic level in the goldap section would favour the s. praecursor zone. in addition, wierzbowski (1978) has described a. lineatum and a. bauhini together from the lower part of the i. planula zone; thus, the a. lineatum subzone seems to correspond to the lower part of the p. baylei zone rather than to the upper part of the r. pseudocordata zone. malinowska (1988) reported specimens of a. bauhini only from the lower kimmeridgian, but these forms belong to other species such as a. bayi or a. cf. cricki. in the subboreal province, the pictonia baylei zone consists of two or three horizons. the lowermost horizon in great britain and the boulonnais area is the pictonia densicostata horizon; as mentioned above, this probably corresponds to the a. bauhini horizon. in the boulonnais and normandy areas, this is followed by the pictonia baylei horizon sensu stricto, and, more widespread in france, the p. baylei and p. thurmanni horizon. in dorset, the second horizon is apparently missing (hantzpergue 1989), while the third one is represented by the p. baylei and p. normandiana horizon, which can also be observed in east greenland (p. aff. normandiana horizon, birkelund & callomon 1985). p. normandiana is regarded as a synonym of p. thurmanni by hantzpergue (1989). this third horizon also contains a. bayi. what conclusions can be made from all these observations? 1. it seems likely that a. bauhini has a longer range in south germany, as suggested by the many records of this ammonite species from the e. hypselum subzone of the e. bimammatum zone to the i. planula and s. galar zones and even from the s. platynota zone; a number of these determinations, although probably not all, may however be erroneous (schweigert 1995b). data from poland also demonstrate that the range of a. bauhini is not restricted to the upper t. hauffianum zone (= a. bauhini horizon), but extends as in southern germany from the upper e. hypselum subzone of the e. bimammatum zone to the top of the i. planula zone sensu lato (matyja & wierzbowski 1997, 1998). it is likely, therefore, that the a. bauhini zone is of longer duration in the submediterranean area, because it comprises not only the a. bauhini horizon of the upper t. hauffianum subzone, but also three or four horizons of the i. planula zone sensu stricto and at least one horizon of the lower s. galar zone. as mentioned above, malinowska (1991, p.16–17) apparently introduced the term a. lineatum subzone for such an extended a. bauhini zone. approximately the same time interval has been called the a. bauhini subzone (of an unnamed zone) by matyja & wierzbowski (1994, 1995) and subsequently elevated to the a. bauhini zone (matyja & wierzbowski 1997, 1998); this zone is now correlated with the upper e. bimammatum zone and the i. planula zone sensu lato. it should also be noted that there is some evidence, at least in scotland, that above the p. densicostatum bed follows another, younger bed with a. bauhini and pictonia sp. (wright 1989). this could be a hint that there are some more beds with a. bauhini, but without p. densicostata, which could correspond to the higher horizons of the p. baylei zone. in england, in contrast, cox & richardson (1982) observed a. bauhini in the uppermost part of the a. rosenkrantzi (= r. pseudocordata) zone. if these determinations are correct, a. bauhini may occur a little earlier than the p. densicostata horizon. one can conclude from these observations that the range of a. bauhini, even in the subboreal regions, is not restricted to the p. densicostata horizon or the ‘a. bauhini zone’ sensu wierzbowski & smelror (1993). 2. if it can be confirmed that amoeboceras bayi and amoeboceras subtilicaelatum are synonymous, as assumed by schweigert (1995b), then the upper horizon of the sutneria galar zone (a. subtilicaelatum horizon) may correspond to the amoeboceras bayi horizon of the lowermost kimmeridgian amoeboceras kitchini zone. it should be noted, however, that a. bayi has also been reported from the lower (‘orthosphinctes’) horizon of the s. platynota zone (atrops et al. 1993b). 3. (a) it can be concluded from the above that correlation of the a. bauhini and p. densicostata horizon with the a. subtilicaelatum and a. bayi horizon is possible, but the vertical ranges of the former species may be longer and the correlation may thus be only partial. consequently, the position of the upper boundary of the a. bauhini zone and the lower boundary of the a. bayi horizon require more precise definition. (b) in the sequence between the a. bauhini and the a. subtilicaelatum horizons, equivalent to the middle part of the pictonia baylei zone, the p. baylei horizon of normandy and the upper a. bauhini-bearing beds in scotland (e.g. bed 38 with pictonia sp., wright 1989) could be expected. they may have 88 their equivalents anywhere in this succession, whereas other parts of the submediterranean succession are not represented in the subboreal sections or only by gaps. (c) the inclusion of this part of the submediterranean subdivision in an a. lineatum subzone (malinowska 1991) with its unprecise limits (in southern germany, the species is known to occur in the upper oxfordian and lower kimmeridgian) will not help significantly; this subzone can be replaced by the a. bauhini zone, as used by matyja & wierzbowski (1997, 1998). (d) there are apparently different possibilities of correlation and further research is necessary to clarify the situation. (e) the subboreal oxfordian–kimmeridgian boundary (r. pseudocordata/p. baylei zone) can, with a high degree of probability, be positioned within the submediterranean and mediterranean scheme in the uppermost part of the e. bimammatum zone on the basis of the correlation of the a. bauhini horizon with the p. densicostata horizon. the submediterranean oxfordian–kimmeridgian boundary remains at the base of the s. platynota zone. additional remarks on lower kimmeridgian correlation as mentioned above, the upper s. galar zone (a. subtilicaelatum horizon) is probably an equivalent of the amoeboceras bayi horizon (schweigert 1995a), which extends into the lower part of the s. platynota zone (amoeboceras horizon with a. bayi, see atrops et al. 1993b). this contrasts somewhat with the correlation of birkelund et al. (1983, table 1), who considered the pictonia baylei zone and the paraspidoceras rupellense zone of hantzpergue (1979) to be equivalent. hantzpergue (1989), too, correlated the p. baylei zone with the p. rupellense zone (horizons r1 and r2); horizons p1–3 of the i. planula zone sensu lato are considered to be equivalent to the r. pseudocordata zone sensu lato (hantzpergue 1989, tables e, f). in his sections, he found the upper oxfordian sutneria galar in the lithacosphinctes gigantoplex horizon (p3), immediately below his p. rupellense zone (see fig. 2). the p. rupellense zone itself is situated between the gigantoplex horizon (p3) of the uppermost idoceras planula zone sensu lato and the rasenia cymodoce zone (fig. 4); it is therefore considered to be equivalent to the lowermost submediterranean kimmeridgian (s. platynota zone; schairer 1970; atrops 1982; olóriz & rodríguez-tovar 1996); its lower horizon (r1) seems to correspond to the upper part of the lower (‘orthosphinctes’) subzone of the s. platynota zone, whereas the lower part (amoeboceras horizon) of this zone is not represented; its upper horizon (r2) contains the index ‘ardescia virgatoides’, which is similar to forms of the ardescia desmoides horizon of the ardescia desmoides subzone of the middle sutneria platynota zone and is therefore very important for correlation to the submediterranean region. above the p. ruppelense zone, hantzpergue (1989) subdivided the rasenia cymodoce zone into nine horizons (c1–9); the r. cymodoce horizon (c2) could be traced from western france to normandy and the subboreal regions. in northern europe, the r. cymodoce horizon is rather widespread (wierzbowski 1989) and in spitsbergen it represents the only rasenoid horizon within the amoeboceras succession. in east greenland, birkelund & callomon (1985, fig. 5) recognised two other horizons below the horizon of rasenia cymodoce (‘17’), namely the ‘pachypictonia’ horizon (‘16’) and the rasenia inconstans horizon (‘15’). these horizons of the lower r. cymodoce zone were considered to be equivalent to the p. altenense horizon (c1; hantzperque 1989); they are probably equivalent to the lower ataxioceras hippolytense subzone of the lower ataxioceras hypselocyclum zone of south-east france, whereas the r. cymodoce horizon perhaps has its equivalents in the upper part of this subzone. in the middle and upper part of the r. cymodoce zone, only a few possibilities remain for far-reaching correlations in europe, such as the eurasenia aulnisa horizon (c5), which contains the highly characteristic submediterranean subzonal index a. lothari, and the semirasenia askepta horizon (c7), which has been found in scotland, england, normandy, western france (birkelund & callomon 1985; hantzpergue 1989) and southern germany (heller 1964; doben & heller 1968). in northern germany, submediterranean ammonites of early kimmeridgian age have been found in sediments which had earlier been attributed to the upper oxfordian (fischer 1991). substages the kimmeridgian stage has been subdivided into two or three substages; here a subdivision into three sub89 stages is preferred. if the middle kimmeridgian is not recognised, then the middle and the upper part are united as upper kimmeridgian (fig. 4). zones in the mediterranean and submediterranean provinces, the lower kimmeridgian consists of three zones, which can be correlated approximately as follows: (1) sowerbyceras silenum – sutneria platynota, (2) ataxioceras hypselocyclum – taramelliceras strombecki and (3) crussoliceras divisum – mesosimoceras herbichi (fig. 4). their further subdivision into subzones is different in both areas (fig. 4); precise correlation of these units is thus difficult (pavia et al. 1987; sarti 1993). detailed subdivisions into subzones and faunal horizons have been proposed in south-east and western france (atrops 1982; hantzpergue 1989); that of south-east france can also be used with some minor changes in southern germany. the submediterranean zonal subdivision as established by geyer (1961) can be used from the iberian peninsula to bulgaria and turkey (sapunov 1977a; lopez marques 1983; alkaya 1992). in poland, the submediterranean zonal subdivision has been adopted by malinowska (1988) and matyja & wierzbowski (1998). in subboreal and boreal regions, subdivision into two zones is typical (see above): (1) pictonia baylei and (2) rasenia cymodoce. these zones can be replaced by the amoeboceras kitchini zone in areas where no perisphinctids occur (e.g. wierzbowski & smelror 1993); this zone may extend into the lower part of the aulacostephanus mutabilis zone (see below). the middle and upper kimmeridgian substages together consist of three zones in all parts of europe (fig. 4). 1. in mediterranean and submediterranean europe: (1) aspidoceras acanthicum zone, (2) mesosimoceras cavouri or aulacostephanus eudoxus zone and (3) hybonoticeras pressulum/h. beckeri or h. beckeri zone. 2. in boreal and subboreal europe: (1) aulacostephanus mutabilis zone, (2) a. eudoxus zone and (3) a. autissiodorensis zone. in regions where no perisphinctids are present, these latter zones can be replaced in the lowermost parts by the amoeboceras kitchini zone (see above) followed by the a. kochi, a. elegans and suboxydiscites taimyrensis zones (fig. 4). the latter index has been taken from northern siberia charts (birkelund & callomon 1985), but there is no mention of this species in more western regions, with the exception of a determination from the middle kimmeridgian of greenland. therefore, for these boreal regions too, aulacostephanus autissiodorensis seems to represent the more appropriate index species. the middle kimmeridgian zonal and subzonal subdivisions can be applied without great difficulty in boreal, subboreal and submediterranean europe, as there are large regions with overlapping guide fossils, whereas in the mediterranean province, only a zonal subdivision is possible. hantzpergue (1989) established a detailed subdivision in western france, which can also be used in northern france (geyssant et al. 1993; proust et al. 1993) and traced as far as germany (zeiss 1991b; schweigert 1993a, 1996a), england, norway and east greenland (hantzpergue 1989). an unresolved problem is the lower boundary of the a. mutabilis zone; it is drawn at the base of the a. lineatum horizon in western and northern france (hantzpergue 1989; hantzpergue et al. 1997), but in england, following the revisions of birkelund et al. (1983), it is placed four horizons deeper, at the base of the s. askepta horizon. recent investigations in central poland came to similar results (matyja & wierzbowski 1998); these workers traced the boundary to a slightly deeper level in the upper a. hypselocylum zone. in germany and the submediterranean region, the usage from south-east france has been followed (hantzpergue 1989; hantzpergue et al. 1991; zeiss 1991b), which facilitates correlation with the base of the a. acanthicum zone; the lower boundary of this zone in germany is traditionally drawn at the incoming of the first representatives of the genus aulacostephanus (lineatum group). in a recent publication by hantzpergue et al. (1997), the problems of this boundary are well illustrated by their table 12; in the ‘biome franco-germanique’, the lower boundary of the a. mutabilis zone is drawn below its lowermost horizon (linealis horizon), whereas the base of the a. mutabilis subzone, curiously, is placed two horizons higher (attenuatus horizon). it is evident that the new data from poland (matyja & wierzbowski 1998), which place the base of the a. mutabilis zone much deeper, will probably necessitate revision of all these correlations. amoeboceras subdivisions are important from norway to spitsbergen (wierzbowski 1989; wierzbowski & århus 1990; wierzbowski & smelror 1993) and east greenland (birkelund & callomon 1985). 90 in the upper kimmeridgian (upper a. autissiodorensis zone) of poland and the russian platform, a sarmatisphinctes fallax subzone has been established (mesezhnikov 1984, 1988; kutek & zeiss 1994, 1997). for the lower part (lower a. autissiodorensis zone), the discosphinctiodes subborealis subzone is proposed; d. subborealis is a significant index fossil. in poland, aulacostephanus autissiodorensis has been found only in the lower and middle parts of the s. fallax subzone. in western siberia, a zone of virgataxioceras dividuum is the equivalent of the s. fallax subzone (mesezhnikov 1988). in northern germany, schweigert (1996a) stated, based on re-study of previous collections, that the a. autissiodorensis zone is probably present. in southern germany, where subdivision into two subzones was previously adopted, new discoveries of ammonites have made it possible to organise the h. beckeri zone into three subzones: (1) sutneria subeumela, (2) virgataxioceras setatum and (3) lithacoceras ulmense (schweigert & zeiss 1994, 1999); further subdivision into several faunal horizons is possible (schweigert 1996b, 1998). furthermore, schweigert (1993a, b, 1994) discovered ammonites in the upper kimmeridgian of swabia with a subboreal habitus, providing better correlation possibilities between the subboreal a. autissiodorensis and submediterranean h. beckeri zones (see below). for the upper kimmeridgian of western france, a useful subdivision has been proposed by hantzpergue (1989), who subdivided the a. autissiodorensis zone into two subzones, the a. autissiodorensis and the gravesia irius subzones, each with two faunal horizons. the succession in the boulonnais area and farther north has been worked out in detail by geyssant et al. (1993) and geyssant (1994); the succession in southern england was reported by cox & gallois (1981), birkelund et al. (1983) and callomon & cope (1996). correlation many difficulties are encountered in correlating zones (and subzones) of the lower kimmeridgian in europe, mainly between the submediterranean and subboreal regions, but also between the submediterranean and mediterranean areas (fig. 4). many correlations are arbitrary and well-constrained correlation is only possible at certain levels. such correlation possibilities in the lower kimmeridgian substage have already been explained in connection with the problems of the oxfordian–kimmeridgian boundary. some problems exist around the lower–middle kimmeridgian boundary, as the base of the a. mutabilis zone is variably defined in different parts of europe (see above). considering the most recent results from poland (matyja & wierzbowski 1998), the lower boundary of the subboreal a. mutabilis zone lies within the uppermost part of the submediterranean a. hypselocyclum zone, i.e. one zone deeper than previously assumed. in the middle kimmeridgian substage, correlations within the a. acanthicum/a. mutabilis zones and the a. eudoxus zone pose no great problems although the uppermost part of the a. eudoxus zone of western france (a. contejeani subzone) seems to correspond to the lower part of the h. beckeri zone in south germany (schweigert 1993b). correlation of the a. kochi zone with the upper part of the a. mutabilis and/or the lower part of the a. eudoxus zone (wierzbowski & smelror 1993) is still tentative, as is the correlation of the a. elegans zone with most of the a. eudoxus zone. correlation of the upper kimmeridgian substage (submediterranean h. beckeri zone with the subboreal a. autissiodorensis zone) was hitherto only possible by indirect arguments. the elaboration of a new zonal and subzonal subdivision in western france by hantzpergue (1989) and the new discoveries by schweigert (1993a, b, 1994) in germany and by kutek & zeiss (1997) in poland now permit correlation of parts of the upper kimmeridgian of western, central and eastern europe and perhaps also western siberia. chronometric data the duration of the kimmeridgian stage has been estimated to be 3.4 ma (gradstein et al. 1995; ogg 1995; ogg & gutowski 1996); for precise data, see figure 4. tithonian and volgian (fig. 5) the tithonian, and its boreal equivalent the volgian, have been confirmed as stage names by a vote of the international subcommission on jurassic stratigraphy in 1990 (zeiss 1991a). a further stage name ‘bononien’ (for the ‘upper kimmeridgian sensu anglico’, proposed by cope 1993) seems unnecessary and could result in each region with a differing zonal subdivision claiming its own stage name, leading only to more confusion rather than to international agreement concerning uniform nomenclature. furthermore, due to the different meanings of the stage ‘portlandian’ in different 91 countries, it was voted in 1990 that usage of this name should be discontinued. the most recent review of the tithonian stage and its ammonites is that provided by geyssant (1997); for the volgian stage and ammonite biostratigraphy, see gerasimov et al. (1995), callomon & cope (1996) and kutek & zeiss (1997). lower boundary the base of the tithonian stage is defined by the base of the hybonoticeras hybonotum zone. it is generally supposed that the base of the coeval gravesia gigas, virgatosphinctoides elegans and ilowaiskya klimovi zones are drawn at approximately the same time level (see also below). substages the tithonian is subdivided into two or three substages; here preference is given to a tripartite tithonian stage (fig. 5). if only two substages are used, then the lower and middle part are united as the lower substage (‘danubian’), the upper substage corresponds to the ‘ardescian’ substage. type regions for the lower and middle tithonian substages have been proposed by barthel (1975) and zeiss (1975). the type region for the upper tithonian substage, the ardescian, has been revised by cecca et al. (1989a, b). the subdivision of the volgian is threefold, into lower, middle and upper substages. the lower and middle substages (‘gorodishchian’) correspond roughly to the tithonian stage (fig. 5), whereas the upper substage 92 n. italy (s. spain) e. austria, moravia russian platforms. germany central poland england greenland vulgaris (durangites) [crassicollaria] subpalmatus ? palatinus vimineus triplicatus tagmersheimense moernsheimense rueppellianus riedense palmatus glaber ciliata rothpletzi/ penicillatum albertinum (darwini) hybonotum volanense (ponti, ‘burckhardticeras’) transitorius transitorius [crassicollaria] [granulosa p.p.] [dunkeri] pseudoscythica pseudoscythica sokolovi sokolovi klimovi klimovi puschi regularis zarajskensis scythicus quenstedti oppressus nikitini blakei rosanovi virgatus zarajskensis pavlovi (disprosopa, contradictionis) oppressus anguiformis kerberus okusensis glaucolithus albani fittoni rotunda pallasioides elegans scitulus eastlecottensis paravirgatus dorsetensis smedmorensis wheatleyensis reisiformis encombensis scitulus wheatleyensis hudlestoni pectinatus primus iatrensis rugosa communis liostraca gracilis pseudaperta anguinus groenlandicus elegans vogulicus (pseudoscythica) mucronatum lithographicum admirandum/ biruncinatum semiforme/ verruciferum richteri richteri volanense austriacus [rugosa] tenuicostata occidentalis tenuicostata tenuicostata magnum [boneti] scruposus simplisphinctes m ic ra ca nt um u pp er m id dl e t ith on ia n lo w er fa lla ux i fa lla ux i sc yt hi cu s pa nd er i pe ct in at us h ud le st on i w he at le ye ns is pa ra vi rg at us v ir ga tu s m id dl e vo lg ia n lo w er v ol gi an n ik iti ni h yb on ot um m uc ro na tu m v im in eu s se m ifo rm e 14 4. 0 (± 2. 5) 15 0. 7 (± 3. 0) mediterranean submediterranean subboreal eastern western boreal fig. 5. a tentative correlation chart for the tithonian and volgian stages in europe (thick lines as in fig. 2). modified after barthel (1964), zeiss (1968, 2001), cope et al. (1980), callomon & birkelund (1982), kutek & zeiss (1988, 1997), mesezhnikov (1988), sarti (1988), zeiss & bachmayer (1989), mitta (1993), kutek (1994) and geyssant (1997). non-ammonite taxa are indicated in square brackets. (‘kashpurian’) belongs to the cretaceous system (sasonova & sasonov 1979; zeiss 1983, 1986; sey & kalacheva 1993a; w.a. wimbledon in: callomon & cope 1996). a type section for the volgian stage has been proposed by gerasimov & mikhailov (1966). zones and subzones whereas the two lower stages of the upper jurassic have two main zonal subdivisions, at least four subdivisions are necessary in the upper stage (fig. 5). this is due to the extreme provincialism of ammonites caused by the increasing isolation of late jurassic marine basins, which seem to have only rarely been directly connected; interbasinal migration was apparently only favoured during the lowermost zone of the stage. the most important lower zone is that of hybonoticeras hybonotum, which can be followed over long distances in mediterranean and submediterranean europe (zeiss 1968; olóriz 1978; sapunov 1979; sarti 1988); in southern germany it is possible to recognise three subzones and seven horizons in the h. hybonotum zone (schweigert & zeiss 1999). in central europe, the latter overlaps with the gravesia gigas zone, which has a rather wide distribution regionally in central and western europe. during the last decades, many new discoveries have been reported and the genus gravesia and the stratigraphy of the beds with gravesia have been revised (hahn 1963; zeiss 1974; hantzpergue 1989; schweigert 1994, 1996a, b; schweigert et al. 1996; zeiss et al. 1996; dimke & zeiss 1997). in the subboreal subprovince, the genus gravesia is also present, but less numerous, so that other index fossils have been given priority, such as virgatosphinctoides elegans in northwestern and ilowaiskya klimovi in eastern europe (cope 1967; cope et al. 1980; kutek & zeiss 1974, 1994, 1997; callomon & birkelund 1982; mesezhnikov 1988). according to callomon & cope (1996), gravesia cf. gravesiana occurs in the lower part of the virgatosphinctoides scitulus zone, thus demonstrating the correlation with the upper h. hybonotum zone (containing g. gravesiana). in northern germany, beds with gravesia gigas intermedia are apparently the youngest beds containing jurassic ammonites (schweigert 1996a) and are succeeded by brackish and freshwater sediments up to the jurassic–cretaceous boundary. in these beds, ostracodes have proved to be the best guide fossil (bischoff & wolburg 1963; schudack 1994, fig. 24), permitting subdivision of the tithonian stage in northwest germany into four zones. in other areas, such as eastern england and denmark, subdivision into nine zones is possible using ostracodes (christensen 1988; schudack 1994, fig. 24). the upper zone of the lower tithonian in mediterranean europe, the zone of semiformiceras darwini (or of virgatosimoceras albertinum), is apparently equivalent to the neochetoceras mucronatum and franconites vimineus zones (each of them with two subzones and some horizons) of submediterranean europe, as they have numerous faunal elements in common (enay & geyssant 1975; olóriz 1978; cecca et al. 1986; sarti 1984, 1988; cecca 1990a, b). precise correlations have still to be worked out, however, and at present this is difficult as no subzones or even horizons have been recognised in the tethyan realm. the submediterranean zones have been traced from south-east france via southern germany to hungary as well as in bulgaria and perhaps also turkey (zeiss 1968; sapunov 1977b, 1979; vigh 1984; fözy 1988, 1993; alkaya 1989; atrops 1994; fözy et al. 1994). correlation with the subboreal regions is only tentative and different proposals have been published (fig. 5; zeiss 1977; mesezhnikov 1988; kutek & zeiss 1997). in subboreal europe, the situation is not much better and correlations between the different subprovinces of northwest and eastern europe are only approximate. consequently, different zonal subdivisions are also applied in these subprovinces. in eastern europe, for example, species of the genus ilowaiskya are used (e.g. the ilowaiskya sokolovi and i. pseudoscythica zones; mesezhnikov 1988; kutek & zeiss 1997), whereas in northwest europe, representatives of the genera virgatosphinctoides, arkellites and pectinatites have been selected (e.g. the virgatosphinctoides scitulus, w. wheatleyensis, arkellites hudlestoni and pectinatites pectinatus zones); each of these latter zones can be subdivided into two subzones (cope et al. 1980; callomon & birkelund 1982; geyssant 1997). for the middle tithonian substage, the subdivisions in mediterranean and submediterranean europe are rather distinct (fig. 5). furthermore, minor faunal differentiations exist within the mediterranean area, and different zonal indexes are used for the same time interval (enay & geyssant 1975; olóriz 1978; cecca & santantonio 1988; sarti 1988): (1) semiformiceras semiforme or haploceras verruciferum, (2) semiformiceras fallauxi or (2a) richteria richteri and (2b) simoceras admirandum/biruncinatum (or s. biruncinatum), and (3) simoceras volanense or ‘burckhardticeras’ peroni or micracanthoceras ponti. note that burckhardticeras olóriz 1978 is a junior homonym of burckhardticeras 93 flores lopez 1967 (schweigert & zeiss 1998). in the submediterranean area of southern germany, the following guiding ammonites have been observed (barthel 1975; zeiss 1986): (1) virgatosimoceras rothpletzi and sublithacoceras penicillatum, (2) lemencia ciliata, (3a) sublithacoceras(?) glaber, (3b) isterites palmatus, and (3c) isterites subpalmatus. according to scherzinger & schweigert (1999), a horizon with sublithacoceras callodiscus has been observed above the level with lemencia ciliata. in eastern europe, the equivalents of the middle tithonian substage are probably the upper part of the lower volgian (upper ilowaiskya pseudoscythica and ilowaiskya tenuicostata zones). the latter unit is discernable in poland but has not been recognised in russia to date (kutek & zeiss 1974, 1988, 1994, 1997; mesezhnikov 1988; kutek 1994). in its upper part, the pseudovirgatites puschi horizon is important due to its mixed fauna (kutek & zeiss 1974, 1988, 1997). a local time equivalent in north-eastern austria is probably the isterites austriacus zone with buchia rugosa as an important guide fossil (fig. 5). a quite different zonal subdivision exists in great britain and the adjoining subboreal and boreal regions as far as greenland (cope 1978, 1980; wimbledon 1980; callomon & birkelund 1982; kejsi et al. 1988); the middle tithonian perhaps corresponds to the main part of the pectinatites pectinatus zone and perhaps to the pavlovia pallasioides zone of england or to the dorsoplanites primus and pavlovia iatrensis zone of east greenland. the upper tithonian substage consists of two or three zones in the mediterranean area. in southern spain, the lowermost zone has been identified as the simplisphinctes zone (tavera 1985). this unit has not been identified in northern italy (sarti 1988), but could be recognised as far as north-eastern austria, where the same ammonite fauna (containing the genus oloriziceras) occurs (zeiss & bachmayer 1989). in the absence of the rather peculiar index genus simplisphinctes, this zone was called the oloriziceras magnum zone for this region (zeiss 2001). above the simplisphinctes (or s. abnormis or o. magnum) zone, the paraulacosphinctes transitorius zone (with the first crassicollaria) occurs. a micracanthoceras micracanthum zone is sometimes adopted instead of the p. transitorius zone; this zone apparently also contains the equivalents of the simplisphinctes (better s. abnormis) zone (enay & geyssant 1975; sarti 1988; geyssant 1997). some authors consider the simplisphinctes and p. transitorius zones as subzones of the m. micracanthum zone (benzaggagh & atrops 1997; geyssant 1997) although the former authors, based on moroccan data, only partially substituted the simplisphinctes subzone, replacing its upper part and the p. transitorius subzone by two new subzones, that of ‘micracanthoceras (corongoceras) spp.’ and that of ‘moravisphinctes spp.’. it is very important that these new subzones can be correlated rather precisely with the calpionellid subdivision; the chitinoidella boneti subzone (of the chitinoidella spp. zone) corresponds to the first two subzones. the base of the crassicollaria spp. zone (zone a) approximately coincides with the base of the moravisphinctes spp. subzone, which corresponds to the lower part of this zone (= subzone a1). the durangites zone follows above the p. transitorius zone. in northern italy, this zone was named the durangites vulgaris zone by sarti (1988); this term has also been adopted by other authors. in some countries, this zone has not been recognised; the equivalents of this zone are then apparently included in the p. transitorius zone, which sometimes even includes parts of the lower cretaceous (e.g. sapunov 1977b). the fauna of this zone has been mainly described by tavera (1985), tavera et al. (1994) and enay et al. (1998a, b). during the middle volgian, central poland belonged to the eastern subboreal subprovince, but only the lowermost unit, the zaraiskites scythicus zone (with the lower z. scythicus and upper z. zarajskensis subzones) is represented (kutek 1994). brackish sediments prevail higher in the polish section and yield ostracodes; the cypridea dunkeri and the cypridea granulosa zones can be recognised. on the russian platform, the lowermost horizon of the z. scythicus subzone (z. quenstedti horizon in poland) is probably represented by beds containing zaraiskites disprosopa and isterites(?) contradictionis (ilovaiskij & florenskij 1941). on the russian platform, a dorsoplanites panderi zone is now used instead of the z. scythicus zone (mesezhnikov 1988; kutek 1994); above follows the virgatites virgatus zone (with three subzones: v. gerasimovi, v. virgatus and c. ivanovi; gerasimov et al. 1995). the v. virgatus zone is succeeded by the epivirgatites nikitini and lomonossovella blakei zone (separated by callomon & birkelund (1982), and, in reverse order, by mesezhnikov (1988) but adopted as a single zone by other russian authors (e.g. gerasimov et al. 1995)). the uppermost middle volgian is represented by the paracraspedites oppressus zone (mesezhnikov 1988). in the baltic area, middle volgian ammonites are rare although a few specimens from lithuania were mentioned by rotkyte. (1976, 1987). in scandinavia, middle volgian ammonites have been found in denmark 94 (birkelund & pedersen 1980) and in norway (birkelund et al. 1978). in england and east greenland, dorsoplanitidae are prevalent, but in both these regions, the subdivisions are distinct; in england, pavlovia pallasioides, pavlovia rotunda and virgatopavlovia fittoni characterise the lower part of the middle volgian whereas progalbanites albani and the giants glaucolithites glaucolithus, galbanites okusensis, kerberites kerberus and titanites anguiformis characterise the upper part (cope 1978; wimbledon & cope 1978). as in russia, the uppermost zone is the paracraspedites oppressus zone (casey 1973; kejsi & mesezhnikov 1986; kejsi et al. 1988), but not all authors adopt this zone. in east greenland, there are some similarities with siberian ammonite successions, but in general the subdivision there has its own character and, with three exceptions, its distinct index species (callomon & birkelund 1982; mesezhnikov 1988): dorsoplanites primus, pavlovia iatrensis, pavlovia rugosa, pavlovia communis and dorsoplanites liostracus characterise the lower part of the middle volgian, whereas dorsoplanites gracilis, epipallasiceras pseudapertum, crendonites anguinus, laugeites groenlandicus and epilaugeites vogulicus are represented in the upper part. the lower part of the upper volgian praechetaites tenuicostatus zone of east greenland may correspond to the uppermost part of the middle volgian, the upper paracraspedites oppressus zone of england and the lower praechetaites exoticus zone (= lowermost craspedites okensis zone sensu lato) of northern siberia. correlation as explained above, the basal zones of the tithonian (and volgian) can be correlated over long distances, but correlation becomes very difficult in the higher parts of these stages. not only is it difficult to correlate between the boreal and mediterranean regions, but also within these regions. distinct lineages of ammonites were evolving throughout the area and consequently it is necessary to develop and apply different ammonite zonal subdivisions; correlation possibilities are thus only few and mostly tentative. many attempts have been made to correlate the different zonal subdivisions of europe (cope & zeiss 1964; zeiss 1965, 1974a, 1979, 1983, 1986; enay 1972; enay & geyssant 1975; olóriz 1978; callomon & birkelund 1982; jeletzky 1984, 1989; tavera 1985; cecca et al. 1986; hoedemaker 1987, 1991; kejsi et al. 1988; kutek & zeiss 1988, 1997; geyssant & enay 1991; sey & kalacheva 1993a; kutek 1994; w.a. wimbledon in: callomon & cope 1996; geyssant 1997). due to problems of provinciality, such correlation schemes are necessarily speculative and ultimately unsatisfactory. a tentative summary correlation scheme is given in figure 5, based on developments since earlier attempts by the author (zeiss 1983, 1986). a similar, although in detail somewhat different, correlation chart has recently been published by hantzpergue et al. (1998). concerning the middle and upper tithonian (upper lower and middle volgian) substages, a number of observations are pertinent. although correlation between the mediterranean and submediterranean area is quite possible in the lowermost middle tithonian substage (s. semiforme/r. richteri – v. rothpletzi/s. pennicilatum zones), a number of different proposals have been made for the higher zones (enay & geyssant 1975; olóriz 1978; jeletzky 1984, 1989; cecca et al. 1986; kutek 1994). a satisfactory answer to this problem requires complete revision of the famous submediterranean neuburg fauna and sections, in which some levels with distinct ammonite faunas have already been recognised by barthel (1964, 1975). in eastern central europe (north-eastern austria, moravia, central and southern poland), some submediterranean and mediterranean ammonites genera of middle and late tithonian age are represented by characteristic forms. they sometimes interfinger with subboreal elements, thus providing good potential for correlation (kutek & wierzbowski 1986; kutek & zeiss 1988, 1997; kutek 1994). the i. tenuicostata and z. scythicus zones of central and southern poland, for example, display interesting forms with affinities to both the submediterranean and subboreal provinces. combined with observations from other localities, this facilitates better correlation between these two regions: (1) the pseuvirgatites puschi horizon of the uppermost ilowaiskya tenuicostata zone contains isterites species described from the higher parts of the neuburg beds, i.e. of late middle tithonian age, and (2) the z. regularis horizon of the lower z. zarajskensis subzone (upper z. scythicus zone) contains pseudovirgatites scruposus and calpionellids indicative of the calpionellid zone a, such that correlation is possible with the lower part of the paraulacosphinctes transitorius zone. in the boreal and subboreal provinces, quite different zonal subdivisions exist, mainly based on different perisphinctid groups, such as the pectinatitinae and dorsoplanitinae in england, denmark, norway and greenland and the ilowaiskyinae, virgatitinae and dorsoplanitinae in poland and russia. the correlation of these zones is rather arbitrary, as demonstrated by callomon & birkelund (1982), mesezhnikov (1988) and 95 w.a. wimbledon (in: callomon & cope 1996), and is based mainly on similar, but non-identical species of dorsoplanitinae. chronometric data the approximate duration of the tithonian has been estimated to be 6.7 ma (gradstein et al. 1995; ogg 1995); for precise data, see figure 5. biochronological importance of nonammonite fossil groups: a review the jurassic system is the classic one for subdivision by ammonites. this fossil group has been used with much success since the pioneering work in the last century by workers such as l. von buch, a. d’orbigny, a. oppel, f.a. quenstedt, k.a. von zittel and s. buckman. indeed, this contribution on the chronological subdivision of the upper jurassic of europe has been compiled primarily using ammonites (see above). however, upper jurassic marine sediments of epicontinental shelves, the habitat of ammonites, are not present everywhere in europe, so that ammonites are not always available. it is often necessary, therefore, to utilise other fossil groups with proven stratigraphic value such as bivalves, brachiopods, foraminifera, ostracodes and distinct plant mega-, microand nannofossil groups. radiolarians, calpionellids, conchostracans, insects and vertebrates should also be added to this list; the first two groups are very useful in pelagic sedimentary basins whereas the last ones are used with much success in the stratigraphic subdivision of continental sediments, such as those of central and eastern asia and of north america. the challenging task of correlating between the different fossil subdivision schemes has been addressed for individual groups (e.g. le hégarat & remane 1968; surlyk & zakharov 1982). multidisciplinary correlation charts, typically for microfossil groups, have only been successfully developed within the last two decades. useful though incomplete examples of such schemes, including microand macrofossils, have recently been published by tavera et al. (1994), r. enay (in: cariou & hantzpergue 1997), gramann et al. (1997) and remane (1997). it remains as one of the more important tasks, however, to establish european multidisciplinary correlation charts that incorporate all fossil groups important for biochronology and also include radiometric ages and palaeomagnetic reversal data. during the editorial work, it was brought to the attention of the author that charts fulfilling many of these expectations have recently been published by hardenbol et al. (1998, charts 6–7); of special interest are the chronometric data for most of the biochronostratigraphic units (see below). invertebrate megafossil groups cephalopods – other than ammonite conchs aptychi in the tethyan regions, aptychi have proven to be a useful addition to ammonites for the subdivision of upper jurassic sediments. following the studies of durand & gąsiorowski (1970) and gąsiorowski (1962, 1985), it is possible to differentiate eleven zones of aptychi using four larger groups of aptychi, the lamellaptychi and laevaptychi and to a lesser degree the laevilamellaptychi and punctaptychi. correlation between aptychi and ammonite zones still poses problems (a. wierzbowski, personal communication 1998). eliás̆ et al. (1996) also used aptychi ranges for biostratigraphy, but without a zonal subdivision; they preferred a multidisciplinary correlation method using the calpionellid subdivision as reference. belemnites the most recent review of this fossil group is that of doyle & bennett (1995) which includes a section on middle and upper jurassic belemnite groups, including those of europe. this publication presents a comprehensive review of the subject, including the work of saks & nalnyaeva (1964, 1966), riegraf (1980, 1981), combémorel & mariotti (1986), and doyle & kelly (1988); a range chart of the most useful taxa for biostratigraphy of the middle and upper jurassic is included by doyle & bennett (1995). the stratigraphic ranges of some more important polish species have been published by pugazewska (1988) and malinowska (1997) and those of sicily by combémorel & mariotti (1990). recently, combémorel (1997) compiled all data available for the tethys and the boreal region of europe and for each of them presented a correlation scheme with the subdivisions based on ammonites and belemnites; see also hardenbol et al. (1998, chart 7). 96 bivalves the most important group of bivalves for biostratigraphic purposes in the upper jurassic of europe is the genus buchia. in the boreal regions of eurasia and north america, it is of particular importance as a supplement to ammonites. the genus has been the focus of many papers in the last decades such as zakharov (1981, 1987, 1990), surlyk & zakharov (1982), jeletzky (1984), kelli (1990), sey & kalacheva (1993b) and sha & fürsich (1994). an interesting interpretation of the different ranges of buchia species in america and eurasia has been presented by hoedemaker (1987). stratigraphic range lists of selected bivalve species from poland have been published by karczewski & pugaczewska (1988) and malinowska (1997). a correlation chart that is mainly based on buchiid bivalves but also includes other bivalve genera (e.g. retroceramus) has been compiled for northern russia and the circum-pacific regions by damborenea et al. (1992). in the upper jurassic, the stratigraphic resolution of bivalve taxa, with the exception of buchiids, seems to be rather limited and/or needs further research (damborena et al. 1992). for some regions, stratigraphic range lists of selected bivalve species have been published, for example for poland (malinowska 1997; karczewski & pugaczewska 1998) and for northern germany by kaever et al. (1976). gastropods the biostratigraphic resolution of this group in the jurassic is not very high, but in special cases, when other guide fossils are not present, some representatives of the group may be used. an example from the upper jurassic of france (nerineaceae) has been published recently by barker (1994). range lists of selected species from poland have been published by karczewski (1988) and malinowska (1997). brachiopods the most recent reviews of this group with respect to upper jurassic brachiopods are those of ager (1994) and alméras et al. (1991, 1994), especially for france and britain, and boullier & laurin (1997) for the tethys and the ‘domaine nw européen français’. ager (1994) considered the group within a global context. alméras et al. (1994) discussed the facies dependence of brachiopods, concluding that distinct zonal species of brachiopods are often necessary for different facies. for biostratigraphical purposes, it is possible to subdivide the upper jurassic of england and north-west france into nine zones and some subunits. the polish species have been figured and described by barczyk (1988); range charts are given in malinowska (1997). prozorovskaja (1993) presented an overview of the brachiopod subdivision of the upper jurassic of the southern part of the former ussr. echinoderms to date, there is no subdivision scheme of the upper jurassic with respect to echinoderms. some genera have biostratigraphic value; saccocoma, for example, has been used in some multidisciplinary schemes. thierry et al. (1997) presented range charts of the upper jurassic regular and irregular echinoid genera and species of france, with the expectation that with detailed research it would be possible to create a subdivision scheme comparable to that developed for the brachiopods of france. corals (scleractinians) this group has poor biostratigraphic resolution. its usefulness for stratigraphic purposes is therefore rather limited, also because of the close dependence of corals on ecological factors (rosendahl 1988). nevertheless, beauvais (1988) subdivided the upper jurassic series (except the lower oxfordian) into six zones based on madreporians (scleractinians). polish species with range charts have been presented by roniewicz & morycowa (1988) and range charts were published by malinowska (1997). sponges this fossil group is poorly suited to regional correlation, but some species may be useful for local subdivision; examples from france have been presented by gaillard (1997). vertebrate megafossils jurassic vertebrate fossils are too scarce to be used as guide fossils. nevertheless, if vertebrate remains are 97 studied thoroughly, they frequently provide valuable biostratigraphic information (e.g. elasmobranchian teeth, gramann et al. 1997). it should be mentioned that the jurassic period in europe saw the early evolution of mammals, the flourishing of the first true birds and the first wave of the acme of the dinosaurs. in other parts of the globe, vertebrates have been used for stratigraphy; in north america, for example, turner & peterson (1998) subdivided the upper jurassic morrison formation into four biozones on the basis of dinosaurs, whereas in china, fish are used for subdivision (chen 1990). invertebrate microfossils foraminifera in the 1950–60s, foraminifera were one of the most important microfossil groups, together with ostracodes, for relative age determinations of marine sediments in boreholes; their importance has decreased in more recent times. studies of foraminifera faunas from outcrops in southern germany were reviewed by groiss (1984). an account of epistominian zonation was given by ascoli (1988), who also presented zonations and correlations between east canadian offshore wells and the east european platform (grigelis & ascoli 1995). foraminifera from northern germany were presented by klingler et al. (1962) and gramann et al. (1997). the guide fossils and characteristic species of the upper jurassic foraminifera of poland have been published by bielecka (1988) and styk (1997), those of the russian platform by a.y. azbel (in: mesezhnikov 1989). foraminifera of sweden were studied by norling (1972) and guy-ohlson & norling (1988). a short compilation of upper jurassic foraminifera in britain has been published by shipp & murray (1981), together with a range chart and figures of index species. the most recent reviews of foraminifera of europe have been compiled by ruget & nicollin (1997) on the small benthic forms, and by bassoulet (1997a) on the large forms; see also hardenbol et al. (1998, chart 7). radiolaria this microfossil group, which has been the subject of much scientific research in recent years in europe, is of particular importance in the tethyan region. a comprehensive monograph was recently published by baumgartner et al. (1995) on the radiolarians of the tethys, including a catalogue of all tethyan species. the biochronological potential for subdividing the upper jurassic series into ‘unitary association zones’ (u.a.z.) is well-demonstrated; there are six such zones covering the whole upper jurassic. they have a duration of between 2–6 ma. this monograph demonstrates the significant advances in research into this group, especially if new quantitative concepts, such as the ‘unitary association zones’, are applied to the biochronological subdivision of the upper jurassic. research into radiolarians and their stratigraphic potential has also been on the increase outside the tethys, as demonstrated by recent publications concerning the submediterranean province (riegraf 1987; kießling 1997; zügel 1997; zügel et al. 1998), and even the subboreal and boreal provinces, including the north sea (dyer & copestake 1989), the russian platform and the barents sea (vishnevskaya 1993, 1997, 1998; kozlova 1994). dyer & copestake (1989) introduced a biozonation based on a succession of ten radiolarian events in the kimmeridgian and tithonian. important attempts are also underway to correlate the new peri-tethyan radiolarian assemblages with different microand macrofossil biozonations (vishnevskaya & de wever 1997); owing to strong provincialism, direct correlation between the peri-tethyan and tethyan zonations is still very difficult, but has been undertaken recently (hardenbol et al. 1998, chart 7). ciliata this group is important only in the tethyan region and the surrounding shelf deposits; the most comprehensive studies of the ciliata in recent years have been published as a result of the sümeg meeting (fülöp 1986; remane et al. 1986). polish forms have been reported by nowak (1988) and those of spain by tavera et al. (1994) and olóriz et al. (1995). remane (1997, 1998) recently published informative reviews of the state-ofthe-art of the group, providing tables which include the stratigraphic succession of calpionellid species and the correlation of calpionellid, nannofossil and ammonite subdivisions with magnetostratigraphic events. nearly simultaneously, blau & grün (1997a, b) and grün & blau (1996, 1997) proposed a revision of the calpionellid zonal and subzonal division. for the tithonian stage, they introduced and formally defined two zones and seven subzones; the duration of zones in the jurassic 98 is less than one million years, that of subzones about 300 000 years. important results from the southern tethyan margin have been contributed by benzaggagh & atrops (1995, 1997). these workers provided precise correlation and species range charts for ammonites and calpionellids for the lower part of the calpionellid succession, which previously was poorly known, and clarified the succession of zones and subzones from the middle tithonian semiformiceras fallauxi/chitinoidella dobeni subzone to the upper tithonian durangites vulgaris/ crassicollaria a3 subzone. an important contribution on the calpionellid faunas of the southern and eastern tethyan region of europe was presented by reháková & michalík (1997); the western carpathians and their foreland in moravia were treated by ̌rehánek (1990) and reháková (1995, 2000). in all these last-mentioned publications, the middle/upper tithonian boundary has apparently been drawn a little too high. following the results of benzaggagh & atrops (1995, 1997), this boundary lies between the dobeni and boneti subzones of the chitinoidella zone and not above this zone. ostracodes this group has a rather high stratigraphic resolution and has therefore been used frequently and successfully for the subdivision of sediments in northern germany, poland, england, the netherlands, the north sea basin, france and russia. in a recent monograph, schudack (1994) revised the ostracodes of the upper jurassic in north-west germany, documenting the correlation possibilities of this group in western, central and northern europe. the upper jurassic of north-west germany was subdivided into nineteen ostracode zones, representing variable durations (0.25–2.5 ma; schudack 1996a; gramann et al. 1997). this study also presents a comprehensive list of all important publications on ostracodes. in northern europe, the papers of herngreen et al. (1988), herngreen & wang (1989) and guy-ohlson & norling (1994) deal with this group in the netherlands and sweden, respectively. in poland, bielecka et al. (1988) treated the group, and range charts have been published by j. szteijn (in: marek & pajchlova 1997); danish faunas were described by christensen (1988). the most recent reviews of european ostracodes are those of bodergat (1997) on marine ostracodes and colin (1997) on non-marine ostracodes; see also hardenbol et al. (1998, chart 7). plant microfossils dinoflagellata dinoflagellate cysts have become a widely used supplement to ammonites and are of particular importance in the subsurface. in a recent study, poulsen (1996) emphasised the important role of dinoflagellates in jurassic stratigraphy while comparing the upper jurassic of denmark and poland. the marine upper jurassic of denmark was divided into seven zones and fifteen subzones whereas that of poland was divided into four zones and twelve subzones (poulsen 1996); the dinoflagellate cyst zonation of the jurassic of subboreal europe is reviewed in poulsen & riding (2003, this volume). in great britain, riding & thomas (1992) have delivered the most recent compilation of dinoflagellates. other important papers are those of sarjeant (1979), riley (1980), riley & fenton (1982) and riding & sarjeant (1984); one concerning russia is that of lentin & vozzhennikova (1990). in the netherlands, herngreen et al. (1988; see also herngreen & wang 1989) presented a report on the stratigraphic bioevents based on the first and last appearance of dinoflagellate cyst species which made possible a subdivision into nine zones. in north-west germany, the oxfordian and kimmeridgian has been subdivided into three dinoflagellate zones and eight subzones (gramann et al. 1997). detailed subdivisions for the boreal and tethyan regions have recently been published by hardenbol et al. (1998, chart 7). calcareous nannofossils (coccoliths, nannolith groups) recent advances in jurassic calcareous nannofossil research have been reviewed by bowen (1996), who dealt with several general aspects of this group, such as evolutionary succession, species diversity and longevity, distribution and provincialism, which are all important when regarding the utility of the group for biostratigraphic purposes. if conditions are favourable, then it is possible to subdivide the upper jurassic into five boreal nannofossil zones (with six subzones) or three submediterranean nannofossil zones (with seven subzones); correlation between these two regions is thus still problematic. the calcareous nannofossil bioevents were recently reviewed by gardin (1997). subdivisions and correlations for the boreal/subboreal 99 and the tethyan/submediterranean provinces can be found in hardenbol et al. (1998, chart 7). charophyaceae this group of calcareous algae has received new impetus with respect to its potential for biostratigraphy. in a recent publication, the results of a local zonal subdivision based on charophytes in the lower saxony basin of north-west germany (schudack 1996b) has been correlated firstly with the new european mesozoic charophyte biozonation (riveline et al. 1996), secondly with the subdivisions of other microfossil groups in north-west germany, such as ostracodes and dinocysts, and thirdly with the old micropalaeontological subdivisions for the upper jurassic (malm) of north-west germany (e.g. klingler et al. 1962; wick & wolburg 1962). in north-west germany, from the upper oxfordian to the top of the tithonian, five charophyte zones are now recognised, whereas in other parts of western europe there are only three (schudack 1991, 1993). the stratigraphic resolution of this group is not very high in the upper jurassic. each biozone represents a duration of between 0.5 and over 2 million years. the charophyaceae of western europe have been revised in detail by schudack (1993), and a useful compilation of all new data in europe has been compiled by riveline et al. (1996); see also hardenbol et al. (1998, chart 7). dasycladaceae this group seems to be only locally important for biostratigraphy (e.g. portugal, italy, dinarids); a short review was presented by bassoulet (1997b). spores and pollen the value of pollen and spore grains for stratigraphic subdivision is not very high in the upper jurassic. the palynostratigraphy of sweden (north-west skåne) was discussed by guy-ohlson & norling (1988) in connection with a study of the microflora of some boreholes. it was revealed that “detailed correlation without the presence of dinoflagellates or other significant taxa appears difficult if not impossible” (guy-ohlson & norling 1988, p. 15). in the central graben of the southern north sea, upper jurassic sediments were subdivided into four zones on the basis of sporomorphs (herngreen et al. 1988; herngreen & wang 1989). in north-west germany, the upper jurassic was divided into four zones using spores and pollen (gramann et al. 1997). the group apparently has its greatest importance at the system boundaries; it has been used successfully at the triassic–jurassic boundary and, to a lesser degree, at the jurassic–cretaceous boundary. magnetostratigraphy this important method has become more directly applicable for stratigraphic purposes in the last few decades, especially when combined with radiometric and biostratigraphic data. some of the more important papers on this topic are: (1) mesozoic in general, harland et al. (1990), gradstein et al. (1995); (2) upper jurassic – lower cretaceous, ogg (1983), ogg et al. (1984), odin et al. (1994); (3) oxfordian, steiner et al. (1985), ogg & steiner (1988a), ogg et al. (1991), ogg & coe (1998); (4) oxfordian – lower kimmeridgian, ogg & gutowski (1996); (5) kimmeridgian–tithonian, ogg et al. (1994), (6) jurassic–cretaceous boundary, ogg et. al. 1984), ogg & lowrie (1986), ogg & steiner (1988b), ogg et al. (1991, 1994). in a recent publication on sequence chronostratigraphy of european mesozoic basins, charts with magnetochronostratigraphic units have been compiled together with sequence chronostratigraphic and biochronostratigraphic data (hardenbol et al. 1998, see below); the time span of the upper jurassic contains polarity chronozones m35 (upper part) – m19. sequence chronostratigraphy sequence stratigraphy is gaining in importance in chronostratigraphic correlation, as illustrated recently by the presentation of a framework for the european mesozoic and cenozoic basins (hardenbol et al. 1998). many data have been used and compiled in charts, two of which are important for the upper jurassic. they demonstrate the sequence chronostratigraphy (sequences, t-r facies cycles, major transgressive–regressive cycles) for the boreal and tethyan realms combined with the ammonite biochronostratigraphy and magnetostratigraphy, plotted against the time scale. the upper jurassic of europe starts in the upper half of the transgressive part of the second major t-r cycle (1st order cycle, named north sea cycle) in the jurassic and ends within the regressive phase of this cycle. a 100 total of 21 sequences (3rd order cycles) have been recognised (ox 0–8, ki 1–7, ti 1–6) and three 2nd order t-r cycles (t8b–r10b) in the boreal area, whereas the number in the tethyan area is somewhat lower. a detailed overview of the north sea cycle in europe (from the north sea to south-east france) has been presented by jacquin et al. (1998); marginal areas have been studied as follows: east greenland (surlyk 1991; 2003, this volume), portugal, lusitanian basin (leinfelder & wilson 1998), portugal and spain, south iberian margin (olóriz et al. 1991), south-east france (jan du chêne et al. 2000), switzerland (gygi et al. 1998), west carpathians (reháková 2000) and russia (sahagian et al. 1996). acknowledgements the author expresses his sincere thanks to jon r. ineson who kindly improved the english text and provided many useful suggestions. two critical readers, beris m. cox and andrzej wierzbowski, contributed much to the advancement of this paper by their useful proposals and comments. many colleagues helped with literature; in particular, i should like to mention jean guex, franç̧ois atrops, elie cariou, raymond enay, reinhart gygi, vassili mitta, zdenek vasicek and andrzej wierzbowski. technical help has been provided by f. boehm, h. forke, w. kießling and e. samankassou at erlangen, and by the staff of the geological survey of denmark and greenland, to whom i express my thanks. references ager, d.v. 1994: brachiopod stratigraphy in the jurassic. in: cariou, e. & hantzpergue, p. 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(ed.): verchnaja jura i graniza ees melovoj sistemoj, 14–29. novosibirsk: izdatelsto <nauka>, sibirskoe otdelenie (in russian). zeiss, a. 1983: zur frage der äquivalenz der stufen tithon/ berrias/wolga/portland in eurasien und amerika. zitteliana 10, 427–438. zeiss, a. 1986: comments on a tentative correlation chart for the most important marine provinces at the jurassic/cretaceous boundary. acta geologica hungarica 29, 27–30. zeiss, a. 1991a: report on the voting about the future usage of the kimmeridgian and tithonian stage names. international subcommission on jurassic stratigraphy newsletter 20, 16 only. zeiss, a. 1991b: ein neuer aspidoceras-fund aus dem oberen jura norddeutschlands und seine bedeutung für die biostratigraphie des norddeutschen kimmeridge. osnabrücker naturwissenschaftliche mitteilungen 17, 87–94. zeiss, a. 2001: die ammonitenfauna der tithonklippen von ernstbrunn, niederösterreich. neue denkschriften des naturhistorischen museums in wien 6, 115 pp. zeiss, a. & bachmayer, f. 1989: zum alter der enstbrunner kalke (tithon; niederösterreich). annalen des naturhistorischen museums wien a 90, 103–109. zeiss, a., schweigert, g. & scherzinger, a. 1996: hegovisphinctes, n. gen., eine neue ammonitengattung aus dem unter-tithonium des nördlichen hegaus und einige bemerkungen zur taxonomie des lithacoceratinae. geologische blätter für nordost-bayern und angrenzende gebiete 46, 127–144. zhamojda, a.i. 1991: postanovlenie o polozhenii kellovejskogo jarusa v obshtshej stratigrafitsheskoj shkale jurskoj sistemy. (resolution on the position of the callovian stage in the entire stratigraphic scale of the jurassic system.) in: postanovlenija mezhvedomstvennogo stratigraficheskogo komiteta i ego postojannych komissij (resolutions of the interdepartmental stratigraphic committee and its permanent commissions). vypusk 25, 23 only. leningrad: vsegei and iggd (in russian). ziegler, b. 1964: das untere kimmeridgien in europa. in: maubeuge, p.l. (ed.): colloque du jurassique à luxembourg 1962, 345–354. luxembourg: l’institut grand-ducal. ziegler, b. 1977: the ‘white’ (upper) jurassic in southern germany. stuttgarter beiträge zur naturkunde serie b 26, 79 pp. zügel, p. 1997: discovery of a radiolarian fauna from the tithonian of the solnhofen area (southern franconian alb, southern germany). paläontologische zeitschrift 71, 197–209. zügel, p., riegraf, w., schweigert, g. & dietl, g. 1998: radiolaria from the nusplingen lithographic limestone (late kimmeridgian, sw-germany). stuttgarter beiträge zur naturkunde serie b 268, 1–43. manuscript received 4 april 1997; revision accepted 14 june 1999. geological survey of denmark and greenland bulletin 7, 2004, p 37-40 37 steam treatment of contaminated soil and aquifer sediment is a promising method of cleaning soil. the treatment is based on steam injection into a water saturated porous aquifer (gudbjerg et al. 2004), by which the heat transfers the contaminants into the vapour phase, allowing entrapment in an active carbon filter connected to a large vacuum suction device. the treatment is effective against several important groundwater contaminants, including pentachlorophenole and perchloroethylene, typically found in association with industrial processes or dry cleaning facilities. furthermore, as an example of removal of non-aqueous phase liquids (napls) large amounts of creosote have been recovered after steam injection in a deep aquifer (kuhlmann 2002; tse & lo 2002). steam treatment is dependent on the complete heating of the soil volume under treatment. the steam has a strongly adverse impact on trees and other plants with deep root systems within the soil, but no other visible effects have been reported. the aim of the activities undertaken during collaborative projects carried out by the geological survey of denmark and greenland (geus) and the danish institute of agricultural sciences (djf) for the danish environmental protection agency and the local authorities in copenhagen (københavns amt) was to establish to what extent the microbial community was affected by the steam treatment of the soil. a few results from the literature indicate that the microbial activity increases in steam treated soil (richardson et al. 2002), probably due to microbial degradation of the soil contaminants in combination with microbial utilisation of heatkilled organisms. it is, however, not known whether this increased microbial activity is associated with the development of pathogenic micro-organisms; these are typically able to grow at higher temperatures than the general microbial community in soil. geological survey of denmark and greenland bulletin 7, 37–40 (2005) © geus, 2005 steam treatment of contaminated groundwater aquifers – development of pathogenic micro-organisms in soil carsten suhr jacobsen, susanne elmholt, carsten bagge jensen, pia bach jakobsen and mikkel bender denmark hedehusene copenhagen 100 km steam steam ground surface ground water level gases and polluted water removed and cleaned up steam steam collection well contaminated soil fig. 1. sketch of steam treatment facility at a strongly contaminated industrial site at hedehusene, west of copenhagen. inset map shows location. 38 the steam treatment in hedehusene hedehusene is situated approximately 25 km to the west of copenhagen, and the contaminated soil and groundwater aquifer here results from various industrial activities primarily carried out between 1920 and 1970. these activities include a dry cleaning facility and several small workshops. from both types of industry, trichloroethylene and tetrachloroethylene are often found as groundwater contaminants. the groundwater aquifer in hedehusene was known to be contaminated with high concentrations of trichloroethylene, which has been a constant hazard to an important drinking water production well downstream from the site. pumping, treating and recycling water at the site over many years had controlled the distribution of the contamination, but the main contamination was still present at very high levels and has become a long-term threat to continued groundwater extraction. the steam treatment in hedehusene was carried out during the winter 2001–2002. wells delivering steam were buried nine metres below the land surface, allowing the transfer of steam below the contamination plume (fig. 1). the steam was pumped continuously for a period of five months, until the temperature reached 90°c. heating of the soil allowed the transfer of the contaminant to the vapour phase which was then trapped in an active carbon filter. heat-tolerant micro-organisms found at the hedehusene site the site was monitored by sampling surface soil and soil from approximately 50 cm depth on 11 september 2001 (before the steam treatment), and resampling during and after the steam treatment. sampling was undertaken six times with the latest sampling on 26 october 2004. in general, it was found that the number of heat-tolerant micro-organisms increased after the heat treatment, and that some of the heat-tolerant micro-organisms could still be found three years after the 2001–2002 steam treatment. heat-tolerant bacteria are defined as able to continue growing at temperatures of 42°c, and heat-tolerant fungi as those able to continue growing at 37°c. such high temperatures do not occur naturally at the site, and soil micro-organisms originating from this site are not expected to be able to grow at such high temperatures. heat tolerance is one of the main characteristics that distinguish normal soil microorganisms from pathogenic micro-organisms found in human patients. general microbial community adaptation to growth at high temperature the effect on the general microbial community was investigated by assessing its growth rate on 24 different microbial food sources. a small amount of soil was added to 24 different microbial growth media and incubated at either 20°c or 42°c. this technique revealed that the microbial community in the control plot was very constant in its ability to utilise the different food sources during the sampling period. furthermore, the microbial community in the control plot showed little ability to utilise food sources at the elevated temperature (an area approximately 30 m away from the heating zone). in contrast, the heated soil showed a massive and long-lasting a b c d 5 cm fig. 2. changes of microbial metabolic fingerprints using comparisons of the ability of micro-organisms to grow on different carbon substrates. the appearance of coloration in each section indicates growth of microorganisms. a high number of positive sections at 42°c indicate a high number of organisms able to grow at temperatures associated with pathogenic micro-organisms. a: steam-treated soil with growth at 20°c. b: control soil with growth at 20°c. c: steam-treated soil with growth at 42°c. d: control soil with growth at 42°c. 39 increase in the ability of the microbial community to utilise the food sources at 42°c (fig. 2). it is well known that micro-organisms differ in their ability to survive in soil. some are able to form spores that can stay inactive in the soil for years while others die out due to predation and competition with micro-organisms having a very low level of metabolic activity. we have chosen two different representatives of heat-tolerant micro-organisms: a bacterium without the ability to form spores, and a fungus which forms conidia. although these conidia are able to germinate and grow in the laboratory, they need not be active in the soil. both species showed a clear response to soil heating as described below. aspergillus fumigatus – an unusual pathogenic and allergenic micro-organism aspergillus fumigatus is a remarkable and unusual pathogen because in addition to causing life-threatening invasive disease of immuno-compromised human patients, it can also cause allergic reactions in persons with fully functional immune systems (latgé 1999; denning et al. 2002). a. fumigatus is easily identified and is distinguished by rapidly growing colonies in characteristic turquoise to dark green colours, by the phialides curving to be roughly parallel to each other and to the axis of the stipe, and the presence of small conidia in columns (fig. 3; klich & pitt 1988). a. fumigatus is regularly reported as a dominant species in various types of compost, but never as a dominant species in soil (domsch et al. 1993). a. fumigatus was only found in very low numbers in the untreated control plot, but in the heat-treated soil this fungal species was abundant. a. fumigatus was still present in elevated numbers at the last sampling in october 2004 in the heat-treated soil, but the numbers were slowly declining. it seems, however, likely that the elevated numbers of a. fumigatus will continue for some time due to the ability of the fungus to form conidia. pseudomonas aeruginosa – an opportunistic pathogen pseudomonas aeruginosa is an opportunistic pathogen, meaning that it exploits any defects in the human host defences to initiate an infection. it causes urinary tract infections, respiratory system infections and also bone and joint infections. furthermore, it is associated with gastrointestinal infections and a variety of systemic infections, particularly in patients with severe burns and in immuno-compromised cancer and aids patients. p. aeruginosa infections are a serious problem for patients hospitalised with cancer, cystic fibrosis and burns. the case fatality rate for these patients is 50%. p. aeruginosa increased from non-detectable (less than 100 cells per gram of soil) in the non-treated soils to 105 cells per gram of soil in the heat-treated soil (fig. 4). p. aeruginosa is a 1 cm fig. 3. colony of aspergillus fumigatus isolated from steam-treated soil. fig. 4. pseudomonas aeruginosa isolated from steam-treated soil. the photograph was taken in ultraviolet light to show the characteristic fluorescence of this bacteria genus. diameter of view is 9 cm. representative of fast growing soil bacteria that are unable to form spores. in contrast to a. fumigatus, the population of p. aeruginosa decreased rapidly after the heat treatment, and after one year the numbers were again below the detection level. this reduction was probably due to predation and lack of competing abilities when the temperature decreased. need for monitoring of microbial sideeffects in relation to steam treatment micro-organisms differ in their ability to develop resting forms. a. fumigatus develops conidia that can remain in soil for many years. these resting conidia may not be active in the soil, even if they can be detected on agar plates when analysed in the laboratory. in contrast, p. aeruginosa does not form resting spores, and detection on agar plates in the laboratory is connected to activity of the bacterium in the soil. the present project highlights the need for microbial risk assessments in connection with new steam treatment projects. the high level of potentially pathogenic micro-organisms expected after heat treatment of a soil points to the need for monitoring these organisms in connection with new steam treatment projects. references denning, d.w., anderson, m.j., turner, g., latgé, j.-p. & bennett, j.w. 2002: sequencing the aspergillus fumigatus genome. the lancet infectious diseases 2, 251–253. domsch, k.h., gams, w. & anderson, t.-h. 1993: compendium of soil fungi 1, 2, 2nd edition, 860 pp., 406 pp. eching: ihw-verlag. gudbjerg, j., trotschler, o., farber, a., sonnenborg, t.o. & jensen, k.h. 2004: on spurious water flow during numerical simulation of steam injection into water-saturated soil. journal of contaminant hydrology 75, 297–318. klich, m.a. & pitt, j.i. 1988: a laboratory guide to the common aspergillus species and their teleomorphs, 116 pp. north ryde, australia: csiro division of food processing. kuhlman, m.i. 2002: analysis of the steam injection at the visalia superfund project with fully compositional nonisothermal finite difference simulations. journal of hazardous materials 92, 1–19. latgé, j.-p. 1999: aspergillus fumigatus and aspergillosis. clinical microbiology reviews 12, 310–350. richardson, r.e., james, c.a., bhupathiraju, v.k. & alvarerez-cohen, l. 2002: microbial activity in soils following steam treatment. biodegradation 13, 285–295. tse, k.k.c. & lo, s.-l. 2002: desorption of pcp-contaminated soil: effect of temperature. water research 36, 284–290. authors’ addresses c.s.j., p.j. & m.b., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: csj@geus.dk s.e., danish institute of agricultural sciences, blichers allé 20, dk8830 tjele, denmark. c.b.j., copenhagen county, soil and groundwater department, stationsparken 27, dk-2600 glostrup, denmark. 40 geologicial survey of denmark and greenland bulletin 2, 96 pp. 1 geological survey of denmark and greenland bulletin 2 · 2003 fish otoliths from the paleocene of denmark werner schwarzhans geological survey of denmark and greenland ministry of the environment 2 geological survey of denmark and greenland bulletin 2 keywords copenhagen, denmark, new taxa, otoliths, paleocene (danian–selandian), teleostei. cover the island of sjælland, denmark, and part of skåne, sweden. localities investigated in this bulletin are shown as are some of the otolith species. design and artist: erik morsing, århus. werner schwarzhans ahrensburger weg 103, d-22359 hamburg, germany e-mail: wwschwarz@aol.com scientific editor of this volume: svend stouge editorial secretary: esben w. glendal referees: dirk nolf, belgium and bettina reichenbacher, germany illustrations: jette halskov and stefan sølberg digital photographic work: benny m. schark graphic production: knud gr@phic consult, odense, denmark printers: schultz grafisk, albertslund, denmark manuscript submitted: 30 june 1997 final version approved: 10 april 2003 printed: 8 september 2003 isbn 87-7871-112-6 geological survey of denmark and greenland bulletin the series geological survey of denmark and greenland bulletin replaces geology of denmark survey bulletin and geology of greenland survey bulletin. citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 2, 94 pp. available from geological survey of denmark and greenland øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk or geografforlaget aps fruerhøjvej 43, dk-5464 brenderup, denmark phone: +45 63 44 16 83, fax: +45 63 44 16 97, e-mail: go@geografforlaget.dk © danmarks og grønlands geologiske undersøgelse (geus), 2003 3 contents abstract . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7 introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9 geological setting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10 danian . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10 selandian . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 11 material and localities . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 11 faunal assemblages . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 12 danian . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 12 selandian . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 12 palaeoecological interpretation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 13 comparison with other paleocene otolith assemblages and palaeogeographic interpretation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 15 danian . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 15 selandian . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16 paleocene/eocene boundary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17 mo-clay (upper thanetian to lower ypresian) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 18 evolutionary interpretation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 20 category 1: persistent taxa . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 20 category 2: extinct early specialised taxa . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 21 category 3: extinct plesiomorphic taxa . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 21 category 4: ‘missing links’ . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 21 stratigraphic significance . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 22 systematic taxonomy . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 22 repository . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25 taxonomic description . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25 order elopiformes greenwood et al. 1966 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 26 family pterothrissidae gill 1893 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 26 genus pteralbula schwarzhans 1981 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 26 pteralbula conchaeformis (koken 1885) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 26 genus genartina frizzell & dante 1965 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 26 genartina hauniensis n. sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 26 order anguilliformes regan 1909 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 29 suborder anguilloidei regan 1909 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 29 family anguillidae rafinesque 1810 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 29 genus anguillidarum semisphaeroides n. sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 29 genus anguillidarum sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 30 suborder congroidei regan 1909 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 30 family congridae kaup 1856 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 30 genus conger oken 1817. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 30 conger illaesus n. sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 32 genus rhechias jordan 1922 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 32 rhechias angulosus n. sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 32 order clupeiformes bleeker 1859 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 33 family clupeidae cuvier 1817 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 33 genus clupeidarum rectiventralis n. sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 33 order siluriformes cuvier 1817 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 35 family ariidae günther 1864 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 35 4 genus arius cuvier & valenciennes 1840 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 35 arius danicus koken 1891 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 35 order salmoniformes bleeker 1859 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 35 suborder salmonoidei bleeker 1859 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 35 family salmonidae rafinesque 1815 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 35 genus salmonidarum acutirostratus n. sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 35 suborder argentinoidei berthelsen 1958 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 36 family argentinidae bonaparte 1838 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 36 genus protoargentinolithus n. gen. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 36 protoargentinolithus balticus (roedel 1930) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 36 protoargentinolithus procerus n. sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 38 genus argentina linnaeus 1758 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 40 argentina erratica (roedel 1930) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 40 argentina longirostris n. sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 42 genus argentinidarum sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 44 order aulopiformes rosen 1973 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 44 family aulopidae cope 1872 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 44 genus aulopus cloquet 1816 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 44 aulopus tortus n. sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 44 family chlorophthalmidae jordan 1923 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 46 genus chlorophthalmus bonaparte 1840 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 46 chlorophthalmus postangulatus nolf & dockery 1993 . . . . . . . . . . . . . . . . . . . . . . . . 46 order myctophiformes regan 1911 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 46 family myctophidae gill 1893 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 46 genus myctophidarum schnetleri n. sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 46 genus myctophidarum sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 48 order percopsiformes berg 1940 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 48 genus ?percopsiformorum enigmaticus n. sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 48 order gadiformes goodrich 1909 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 50 family ranicipitidae markle 1989 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 51 genus raniceps oken 1817. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 51 raniceps hermani nolf 1978 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 51 family merlucciidae gill 1884 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 52 genus palaeogadus rath 1859 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 52 palaeogadus sinangulatus n. sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 52 family lotidae bonaparte 1832 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 54 genus gadomorpholithus n. gen. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 54 gadomorpholithus ponderosus (koken 1885) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 54 genus molva lesueur 1819 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 56 molva palaeomorpha n. sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 56 family gadidae rafinesque 1810 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57 genus protocolliolus gaemers 1976 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57 protocolliolus amorphus n. sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57 family macrouridae jordan & evermann 1898 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 59 genus coelorhynchus giorna 1809 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 59 coelorhynchus balticus (koken 1885) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 59 genus coryphaenoides gunnerus 1765 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 61 coryphaenoides amager n. sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 61 genus hymenocephalus giglioli 1884 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 63 hymenocephalus rosenkrantzi n. sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 63 gadiformes spp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 65 5 order ophidiiformes berg 1937 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 65 family ophidiidae rafinesque 1810 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 65 genus ophidiidarum seelandicus (koken 1885) . . . . . . . . . . . . . . . . . . . . . . . . . . . . 65 family bythitidae gill 1861 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 67 genus bidenichthys barnard 1934 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 67 bidenichthys lapierrei (nolf 1978) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 67 genus bythitidarum rasmussenae n. sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 69 order lampridiformes regan 1909 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 70 family veliferidae bleeker 1859 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 70 genus veliferidarum harderi n. sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 70 order zeiformes regan 1909 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 72 genus zeiformorum janni n. sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 72 order beryciformes regan 1909 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 75 suborder berycoidei regan 1909 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 75 family berycidae lowe 1843 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 75 genus hoplostethus cuvier 1829 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 75 hoplostethus lacinatus koken 1885 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 75 genus centroberyx gill 1862 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 77 centroberyx integer (koken 1885) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 77 centroberyx fragilis n. sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 78 order scorpaeniformes garman 1899 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 80 suborder scorpaenoidei garman 1899 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 80 family scorpaenidae risso 1827 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 80 genus scorpaena linnaeus 1758 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 80 scorpaena corallophilus n. sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 80 order perciformes bleeker 1859 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 83 suborder percoidei bleeker 1859 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 83 family apogonidae jordan & gilbert 1882 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 83 genus apogonidarum sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 83 family acropomatidae gill 1893 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 83 genus acropoma guenther 1859 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 83 acropoma sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 83 family carangidae rafinesque 1815 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 85 genus carangidarum sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 85 family sparidae bonaparte 1832 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 85 genus sparidarum sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 85 suborder scombroidei bleeker 1859 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 86 family gempylidae gill 1862 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 86 genus gempylidarum merus n. sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 86 suborder stromateoidei regan 1909 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 86 family centrolophidae regan 1909 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 86 genus mupus cocco 1840 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 86 mupus sinuosus (stinton 1965) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 88 genus stromateoidarum sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 88 order tetraodontiformes berg 1940 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 88 family ostraciidae rafinesque 1815 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 88 genus ostracion linnaeus 1758 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 88 ostracion pergravis n. sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 88 lapilli . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 90 acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 90 references . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 91 6 7 abstract schwarzhans, w. 2003: fish otoliths from the paleocene of denmark. geological survey of denmark and greenland bulletin 2, 94 pp. fish otoliths are described from the lower paleocene (danian) and middle paleocene (selandian) from sjælland in denmark. a total of 44 species are described, 23 as newly established and nine in open nomenclature. thirteen species (including seven new species) have been obtained from the danian poorly consolidated coral limestone at fakse and 39 species (including 19 new species) from the selandian at localities near copenhagen. both stages have previously been poorly known for otoliths in the north sea basin, and are described from only two previous publications, i.e. koken in 1885 from the selandian of copenhagen and roedel in 1930 who described otoliths from erratic ice age boulders in north-east germany. the original material of both workers has been revised in this bulletin. otoliths are well known elsewhere in the north sea basin since upper paleocene (thanetian) times and are described from the london basin and from belgium. palaeographic, palaeoecological and biostratigraphic implications of the otolith findings in the paleocene of the north sea basin are discussed. two new genera and 23 new species are introduced and described. the new taxa are: genartina hauniensis n. sp., genus anguillidarum semisphaeroides n. sp., conger illaesus n. sp., rhechias angulosus n. sp., genus clupeidarum rectiventralis n. sp., genus salmonidarum acutirostratus n. sp., protargentinolithus procerus n. sp., argentina longistrostris n. sp., aulopus tortusn. sp., genus myctophidarum schnetleri n. sp., genus ?percopsiformorum enigmaticus n. sp., palaeogadus sinangulatus n. sp., molvia palaeomorpha n. sp., protocolliolus amorphus n. sp., coryphaenoides amager n. sp., hymenocephalus rosenkrantzi n. sp., genus bythitidarm rasmussenae n. sp., genus veliferidarum harderi n. sp., genus zeiformorum janni n. sp., centroberyx fragilis n. sp., scorpaena corallophilus n. sp., genus gempylidarum merus n. sp. and ostracion pergravis n. sp. author’s address ahrensburger weg 103, d-22359 hamburg, germany. e-mail: wwschwarz@aol.com 8 fig. 1. map of denmark and southern sweden. the location of fakse on the island of sjælland is shown. inset map shows locations in the copenhagen area. 100 m ★ ★ ★ ★ ★ ★ ska ger rak baltic sea 10° 10° 14° 58° 58° denmark amager vestre gasværk københavn norway sweden 100 km n københavn jylland germany odense århus kongedyb ii i gemmas allé lufthavn sundkrogen ålborg fakse malmø fyn sjælland skåne bornholm 9 introduction paleocene otoliths are little known from the pre-thanetian strata. koken (1885) and roedel (1930) described a few species respectively from the copenhagen region and from erratic ice age boulders of north-east germany. both faunas are referred to the selandian stage of the paleocene series and apparently represent similar faunas. the paleocene otoliths described in this bulletin originate from five different localities in denmark (fig. 1). the danian otoliths have been collected from poorly consolidated coral limestone pockets at fakse – the classical danian locality – on the south-east coast of sjælland (fig. 1).this small fauna is remarkable in several aspects: (1) it represents the first otolith collection described from the danian, (2) it is the first otoliths representing a fossil fish fauna from a coral mound environment and (3) the fakse otoliths are amongst the first known otoliths, which have been completely altered from the original aragonite composition to calcite without any obvious loss of morphology. most of the selandian material described here is from the copenhagen area and was collected by p. harder (see harder 1922) and a. rosenkrantz (see rosenkrantz 1930 and ravn 1939). four faunas of selandian age come from localities in the vicinity of copenhagen and the same otoliths previously described by koken (1885) are recorded from here. koken (1885) originally described seven species, however the fauna has now risen to 39 species. roedel (1930) reported 21 species from the paleocene of north-east germany, including 14 as new, but after revision of his material only eight are considered valid (including four species previously described by koken). for a long time roedel’s material was considered to be lost after the second world war, but recently it was found by w.-d. heinrich in the collections of the humboldt-university (berlin). recently it has been discussed whether the traditional three-fold division of the paleocene series should be applied (i.e. danian, selandian and thanetian stages) or if a two-fold division of the paleocene series would be more appropriate (schmitz 1994). in the latter case, the selandian stage should be included the thanetian stage. however, it is formally decided to maintain the three-fold division of the paleocene series (berggren 1994; schmitz 1994) and is followed here. the paleocene mollusc fauna supports a three fold division (schnetler 2001) and the otoliths from the selandian (sensu stricto) strata of the type area and described here also show a number of specific differences when compared to the younger thanetian (sensu stricto) strata from england (stinton 1965). 10 geological setting the danian stage represents the latest stage of chalk deposition in thenorth seabasin thatprevailed throughout the upper cretaceous. theselandiansediments indenmarkcover aneroded danian limestone relief (fig. 2; thomsen 1994). the erosional surface reflects a regional lowstand in the north sea basin. the low-stand is related the onset of the laramide tectonic phase resulting in marginal uplift and basinal subsidence (rifting) just prior to the initiation of the late paleocene sea-floor spreading in the norwegian–greenland sea (ziegler 1982; berggren 1994). with the onset of the selandian transgression a widespread clastic environment became established in the north sea basin, replacing the previously dominating carbonate environment in the late cretaceous and danian times. danian the fakse quarry on sjælland and together with stevns klint are the type areas for the danian stage (desor 1846). the danian limestone is lithological relatively uniform and the unit is referred to the bryozoan limestone and the overlying københavn limestone (stenestad 1976; surlyk & håkansson 1999). the fakse quarry age lithology hiatus hiatus formation holmehus fm danian limestone æbelø fm kerteminde marl lellinge greensand ølst fm fur fm (mo-clay) foram zones nanno zones p6 pa le oc en e se la nd ia n d an ia n t ha ne tia n y pr es ia n eo ce ne p5 np10 np9 np8 np7 np6 np5 np4 np3 np2 np1 p4 p3 p1 hiatus p2 p0 a c b a b a c b a claysandstone limestone marl fig. 2. paleocene stratigraphy. modified from thomsen (1994, 1995). 11 is one of five coral limestone localities known from the danish–polish trough. the stratigraphic position of the fakse quarry is middle danian (biozone np3 = d6–7; fig. 2; thomsen 1995). the danian sediments at fakse represent a carbonate facies consisting of coral bioherms (bernecker & weidlich 1990; surlyk 1997). the buildups are moundor band-like and they are made up of bryozoan and predominantly dendroid coral limestones. laterally the mounds are associated with pelagic limestone facies containing globigerinids and coccoliths (chalk facies). a particular type of biogenic mounds, which is composed of scleractinian corals and sponges, dominates the mounds at the fakse quarry (willumsen 1995a, b). bernecker & weidlich (1990) and surlyk (1997) interpreted the coral limestone as deeper-water coral bioherms that grew at a palaeodepths between 100 and 300 m. according to bernecker & weidlich (1990), the solution of aragonite and neomorphism of calcite most likely indicates diagenesis under meteoric-phreatic conditions, which prevailed during the regression at the end of the danian. selandian the selandian deposits in denmark consist of clastic and fine clastic sediments. the selandian sediments are divided into four formations. the basal transgressive sediments are represented by the lellinge greensand (biozones p3/np4–np5; fig. 2), which are partly overlain by and partly lateral equivalent to the kerteminde marl (biozones p3/np4–np5; fig. 2). the kerteminde marl is succeeded by the non-calcareous æbelø formation (biozones p3/np5; fig. 2). the selandian succession is conformably overlain by the clays of the holmehus formation (biozones p4/np6–np8; fig. 2), which is mostly thanetian age (thomsen 1994; schnetler 2001). the selandian mollusc faunas at copenhagen indicate that relative warmer water conditions prevailed in the danian at fakse than at copenhagen during the selandian time (schnetler 2001). material and localities all the described otoliths originated from fakse quarry, sjælland and from four locations near copenhagen (fig. 1). fakse quarry. preservation of primarily aragonitic otoliths is very rare in the danian limestone. the otoliths described here however have been extracted from poorly lithified pockets that occur in the coral mounds at fakse. vestre gasværk. the largest collection comprising about half of total material originates from vestre gasværk (vesterbro district; fig. 1), which is within copenhagen city. the material was collected by a. rosenkrantz (rosenkrantz 1930) and from sediments that were temporary exposed in the excavation in 1930 for a (former) coal-gasification factory. sundkrogen. a large collection of otoliths is recovered from the sundkrogen. sundkrogen was an excavation made during the deepening of the sundkrogen harbour basin or the orientbassinet in the harbour of copenhagen (district østerbro; schnetler 2001; fig. 1). p. harder and a. rosenkrantz collected the material (harder 1922; rosenkrantz 1920). the mollusc fauna from these two collections has been investigated and described by ravn (1939) and schnetler (2001). schnetler (2001, p. 18) noted that the wells and the exact depths were not annotated by poul harder on his labels. the harder collection has been kept in different drawers and a colour code was used to indicate depth intervals. schnetler (2001) was able to solve the code and relate the samples to certain depths and intervals of the wells and the rocks exposed in the orientbassinet. all the investigated samples are from the lellinge greensand. schnetler (2001) discovered that in the harder’s collection the drawers 1–5 represent the lower part of the lellinge greensand and the drawers 7–9 the upper part of the lellinge greensand. the drawer 14 12 represents a transitional zone composed of fine clayey sand situated between the greensand and the overlaying clay, and drawer 18 contains material composed of dark brown sticky clay. kongedyb i and ii. a small collection originates from the two shallow wells kongedyb i and ii (fig. 1; see ravn 1939). these were drilled in 1934 in øresund and just east of amager. the kongedyb fauna is interesting as it shows significant differences to the two previously mentioned ones. a. rosenkrantz collected the material investigated here. gemmas allé. a small collection has been obtained from the gemmas allé section on amager (fig. 1) and the material was collected by m. nielsen. the lithostratigraphic position of this temporary exposed and now covered locality extended from the top of the danian limestone and into the lellinge greensand (stouge et al. 2000). faunal assemblages danian the danian otolith assemblage from fakse is of low diversity and nearly 90% of the total assemblage are composed of three dominant species. these are the berycid centroberyx fragilis (40%), the bythitid bidenichthys lapierrei (32%) and the scorpaenid scorpaena corallophilus (16%). two of these three species are also known from the younger selandian strata, namely centroberyx fragilis and much less commonly, bidenichthys lapierrei. scorpaena corallophilus is missing entirely from the selandian collection. in total, thirteen otolith species have been identified in the collection from the danian at fakse (excluding larval gadiform specimens, which cannot be identified). seven of these have also been found in the selandian strata, but five species so far are restricted to the danian, i.e. aulopus tortus, genus bythitidarum rasmussenae, scorpaena corallophilus, genus apogonidarum sp. and genus gempylidarum merus. the most likely explanation for this low level of diversity and the apparent difference to the selandian otolith assemblages may be seen primarily in the differences of the environment and facies. selandian the otolith assemblages from the four selandian localities exhibit a high degree of diversity. this is measured by summing the most common species up to the 90% level of the entire assemblage. following this procedure, the range of diversity is between 10 and 15 species.dominant groups arevarious gadiforms, argentinids, ophidiids and pterothrissids. gadiforms and argentinidsare thebest represented groups with respectively eight and five species. pterothrissids are represented by pterothrissus conchaeformis and the problematical genartina hauniensis, and the species genus ophidiidarum seelandicus accounts for virtually all ophidiids. the latter is the most common species in the selandian of denmark when the whole record is considered. the next most common species are the argentinid protargentinolithus balticus, the pterothrissid pterothrissus conchaeformis and the macrourid coelorhynchus balticus. the statistics are somewhat distorted regarding the latter. it must be assumed that a large number of the larval, non-identifiable gadiform otoliths, which form the single biggest element, in fact belong mostly to coelorhynchus balticus. if so thiswould make this taxon the most common species in the selandian of denmark. although variable in the ranking order, these four species and the larval gadiforms are always among the five most abundant in each of the selandian locations, regardless of whether it is a sandy or silty facies. furthermore at the two locations from which stratigraphical collections exist, no major changes occur in the faunal assemblage. argentinid and larval gadiform otoliths seem to be somewhat under represented at gemmas allé and vestre gasværk, but this could be the result of a different collecting technique and thus may not reflect a true change in faunal composition. there are some subtle differences, which more likely represent true differences in the original faunal distri13 bution. for instance, at gemmas allé and vestre gasværk the larger andbetter identifiable gadiform otoliths are more frequent than at sundkrogen. sundkrogen fauna is richer in genartina and argentinid otoliths than at the other locations, whereas ariid, chlorophthalmid and ophidiid otoliths are particularly common at the vestre gasværk locality. kongedyb is remarkable in that the macrourid hymenocephalus rosenkrantzi, which either is very rare or absent from the other locations, is about equally common as the macrourid coelorhynchus balticus. this could indicate that a somewhat deeper depositional environment prevailed at the kongedyb location than at the other locations. palaeoecological interpretation the four main fish groups represented by otoliths in the danish selandian warrant a closer look on their possible palaeoecological preferences. the recent known representatives of pterothrissidae are two endemic species living on the deeper shelf and the upper continental slope of the north-west african and the japanese coasts. this distribution pattern reflects a typical secondary endemism, i.e. a relict of a formerly much wider distribution. in the fossil record their otoliths exhibit an almost worldwide distribution – particularly in the late cretaceous and the palaeogene – but only represented by a few species at that time. at a given fossil location pterothrissidae are usually represented by one or at most two species. occasionally, they can be relatively frequent, as is the case in the selandian deposits of denmark. in the late oligocene of germany pterothrissids are locally common. müller (in schwarzhans 1994) reported on mass occurrences of juvenile pterothrissid otoliths in a very shallow, near shore bay environment. he explained the abundance as the bay environment functioned as a breeding place for these fishes. elsewhere in early tertiary both juvenile and adult pterothrissid otoliths are common in sediments representing shallow-water environment. a similar situation is now observed in the selandian of denmark, where mostly small pterothrissid otoliths form part of the dominant faunal element among the teleosts. clearly, the shallow-water occurrence of pterothrissid otoliths in the palaeogene sediments is not in accordance with the present-day environmental distribution of the members of this family. it is concluded that by analogy with other fossil findings early palaeogene pterothrissids were distributed (almost) worldwide and lived in shallower shelf environments than nowadays. recent argentinidae are typically found schooling close to bottom on the outer continental shelf and the upper slope and with an anti-tropical distribution pattern. some genera such as nansenia and microstoma are mesopelagic and the related family bathylagidae is bathypelagic. in the fossil record argentinid otoliths and related forms occasionally occur in large numbers in upper cretaceous and lower palaeogene sediments, but they have rarely been reported as dominant and neverwith so many species as in the selandian of denmark. argentina erratica is one of the common argentinid species in the selandian strata of denmark. it is also a dominant species in the paleocene (early eocene) of the ellesmere island (schwarzhans 1986; reported as a. pennata (stinton 1966) – see synonymy for a. erratica) and is well known (under a number of synonymies) from the thanetian of the london basin (stinton 1965, 1966). from these observations it may be concluded that the paleocene argentinids from denmark have lived in schools just like their recent counterparts, though probably in shallower water. gadiform otoliths are the dominant and most diverse element inpracticallyallotolithassemblages in sediments of the north sea basin from the middle oligocene to recent times. before that, in sediments of the early oligocene and eocene they are extremely rare. it seems that they have been replacing ophidiiforms, which form the dominant and most diverse element in otolith assemblages in eocene sediments. the gadiform otolith assemblage in the selandian of denmark is among the earliest in the fossil record and certainly the richest in pre-oligocene times. it thus has interesting implications not only for the otolith palaeoecology but also for the palaeogeographic and evolutionary interpretations. furthermore, gadiform otoliths are not uncom14 mon in the upper paleocene – thanetian sediments of the london basin (stinton 1965, 1966) and belgium (nolf 1978). in contrast to the younger gadiform assemblages, which are mainly composed of gadids and merlucciids, the diversity of the selandian gadiform assemblage is also remarkable since it includes common macrourids and lotids. there are four common and four rare specieson the record.the common species are: coelorhynchus balticus (macrouridae), palaeogadus sinangulatus (merlucciidae), protocolliolus amorphus (gadidae) and gadomorpholithus ponderosus (lotidae). in the recentmost fishes in the families gadidae, lotidae and merlucciidae live in schools or individuals on and over the shallow to middle shelf soft-bottom environments. gadidae and lotidae are typical fishes of the cool and temperate seas of the northern hemisphere. merlucciidae are more widely distributed along subtropical and tropical coasts. macrouridae in contrast are benthopelagic fishes in the deep seas and on the continental slopes with some species distribution patterns extending onto the outer, deeper shelf. in this respect it has to be noted that the macrourid coelorhynchus balticus is the single most common species in the selandian of denmark – provided that most of the unidentifiable larval gadiform otoliths represent this species. the abundance and richness of gadiforms in the selandian of denmark results in some controversial conclusions. in comparison with younger assemblages it would call for a temperate, shallow-marine, clastic environment. palaeobathymetry of recent macrouridae is in conflict with the mollusc findings in the selandian (schnetler 2001). the solution to the conflicting interpretation is to assume that in this early time certain ‘primitive’ macrourids were adapted to shallower shelf environments. as already mentioned above, ophidiid otoliths form the dominant and, in terms of species, most diverse element in practically all otolith assemblages of early oligocene and eocene times. in the north sea basin they apparently occupy the ecological space that later becomes the domain of the gadiformes. both groups seem to replace each other (schwarzhans 1981c). in the recent ophidiiforms live benthopelagic and are not quite as common. there are few genera and species, which occur in a variety of shallow tropical environments, but the large part of today’s ophidiids, are deep-water fishes. however, it has long been recognised that the abundant late cretaceous and early tertiary ophidiid otoliths largely represent extinct genera that are related to the nearshore living genera. also, at that time ophidiid fishes formed a much more important and rich component of the warm shallow water and marine teleost fauna (nolf 1980, 1985; schwarzhans 1981c). in contrast to this general observation the selandian assemblages of denmark are very poor in the number of ophidiiform species and the whole group being virtually represented by a single species, i.e. ophidypterus seelandicus. despite the lack of diversity this single species is among the three most abundant species. in conclusion ophidiid otoliths are common in the selandian sediments of denmark (as would be expected in lower tertiary sediments) but represented by only one species. the total lack of species diversity suggests that it was not a very suitable environment for fishes of this family possibly due to unfavourable temperature. following the regional lowstand and erosive stage at the end of the danian caused by marginal uplifting during the laramide tectonic phase (ziegler 1982, 1990), the transgressive selandian sediments in denmark were deposited in a relatively more shallow shelf environment than the danian sediments. the selandian palaeobathymetry in locations near copenhagen was probably about 50 to 100 m according to the analysis of the mollusc faunal association (schnetler 2001) and of the benthonic foraminifera (larsen & jørgensen 1977). at the same time basinal subsidence was initiated in the central north sea basin some 300–400 km away from these locations. in this area water depths was roughly 500 to 900 m (ziegler 1982). the composition of the fish fauna in the danish selandian as reconstructed from the otoliths does not reflect a typical shallow to middle shelf environment. as mentioned above, the otolith association contains a number of elements, which in the recent are typically found on the continental rise and the deeper shelf at water depths of 200 to 500 m. however, a direct comparison between teleosts from the recent and selandian may not always be appropriate for reasons explained above. it is possible that certain fish taxa, which in the recent are adapted to deeper water, previously lived in more shallow-water environments. on the other hand, it is also possible that, in the case of the selandian of copenhagen, deeper water fishes from the subsiding deep-water provinces in the central part of the north sea basin have been transported and deposited by some allogenous mechanism. this shows that early teleost faunal associations, such as in the selandian of denmark, must be treated carefully for palaeoecological and related purposes. in my opinion it is in such in15 stances more meaningful to reconstruct the ways and changes in the living of teleosts through their evolution by using analyses from more reliable fossils such as molluscs or benthic foraminifera. measurements of oxygen isotope ratios from shell material of the various tertiary strata of the north sea basin have been used to interpret the palaeotemperature development in the southern part of the basin (buchardt 1977, 1978). from a slight temperature minimum during danian palaeotemperatures are interpreted to have raised during selandian and thanetian to a warm temperate to subtropical climate. during eocene times and in particular during the middle eocene a temperature maximum is recorded with a warm subtropical to tropical climate. in oligocene temperatures declined rapidly and with some fluctuations and minor peaks in middle miocene and pliocene remained in a temperate to cooler subtropical zone. buchardt (1977, 1978), however, specifically pointed out that measurements were based on shell remains of benthonic organisms and therefore reflect bottom temperatures rather than surface temperatures. this could have had influence on the danian palaeotemperature interpretation if material was obtained from cooler deeper water sediments such as the from the fakse location (k.i. schnetler, personal communication 2000). as stated above, the selandian otolith association from denmark (and to a lesser extend the thanetian one from the london basin as well) contains a large number of teleosts, which in the recent are typical for temperate seas – mostly gadiforms and argentinids. fish groups characteristic for warm subtropical or tropical seas are rare. such groups dominate the eocene otolith associations for instance of belgium and france. there, typical northern hemisphere temperate groups like the gadidae and argentinidae are missing almost entirely. later, frommiddleoligoceneonwardsgadiform dominated otolith associations prevail again throughout the deposits in the north sea basin, indicating temperate to subtropical climatic conditions (for a more detailed discussion see schwarzhans 1994). it thus seems that the otolith findings support buchardt’s palaeotemperature interpretations. comparison with other paleocene otolith assemblages and palaeogeographic interpretation the geological history of the north sea basin during the tertiary is characterised by phases of nearly complete geographic isolation alternating with phases of changing connections with the neighbouring seas (fig. 3). like any other faunal element, the fish fauna has reacted to these changes and this is reflected in sudden changes of the faunal composition. major faunal changes in the otolith composition are observed at the boundaries of the early to middle paleocene, paleocene to eocene, eocene through early oligocene to middle oligocene and less drastic changes in the middle miocene, pliocene and pleistocene. danian descriptions of late cretaceous to early paleocene otolith assemblages are few.there are reports from the late cretaceous of north america (nolf & dockery 1990; nolf & stringer 1996), the north sea basin from northern germany (voigt 1926) and mostly unpublished data from the bavarian basin in southern germany (koken 1891) and a synthesis of data (fig. 4; schwarzhans 1996). these faunas are dominated by beryciformes, elopiformes and anguilliformes. ophidiiformes and chlorophthalmidae are also common. the few danian otolith assemblages known from north america (nolf & dockery 1993) and europe (this bulletin and an unpublished assemblage from bavaria; fig. 4) differ from the late cretaceous ones mainly in the way that the beryciformes have already become much more rareandareprogressivelybeingreplacedbyperciformes. most otoliths of these early perciforms are very primitive and generalised in appearance and thus often difficult to be defined taxonomically. 16 selandian the danian–selandian boundary is marked by important geological events in the north sea basin that also find their reflection in the development of the fish fauna as represented by otoliths. first, the marginal uplifting during the laramide tectonic phase (ziegler 1982) separated the north sea basin from the atlantic, prior to the deposition of the transgressive selandian sediments in denmark (figs 3, 5). this probably influenced the water currents by establishing a counter-clockwise longshore drift that would bring water from higher latitudes of the north sea basin towards the south. possible connections in the north to the kara sea (off western siberia) or via the beginning greenland–norwegian rifting to the arctic ocean may have first occurred during middle to late paleocene (ziegler 1988). the carbonate environment of deposition of the danian age has changed into a terrigenous sequence. the rise of colder water fish groups (gadiformes, argentinidae) and the suppression of the warm water loving ophidiiformes and beryciformes/ perciformes is a clear response to these palaeogeographic, environmental, and climatic changes. during the selandian–thanetian time interval, the otolith assemblages of the north sea basin show some regional differences (fig. 5). four out of eight identified species from the selandian (d. nolf, personal communication 1999) of belgium are also known from the selandianofdenmark. in theenglishthanetian the situation is similar, but stinton’s identifications (1965) need detailed revision. theoverall faunal composition (gadiformes dominating ophidiiformes) is similar. an undescribed middle to late paleocene otolith assemblage from bavaria reflects quite a different situation. besides the fact that the bavarian otolith assemblage contains a number of deep water fish representatives (stomiiformes), it is much richer in warm-water fish groups (ophidiiformes, beryciformes, perciformes) and less rich in cold-water groups (no argentinidae, only rare gadiformes). this calls for an effective separation of the two seas (see ziegler 1988) with different climatic conditions and differences in the geographic origin of the respective faunas due to the palaeogeographic situation. still another undescribed fauna from the paleocene of west greenland shows a much closer resemblance with the danish selandian fauna, even at the species level. these few otolith assemblages known so far from arctic basin la bra do r s ea west siberian basinbarents shelf north atlantic molasse basin c aspian basinn or th s ea b as in fig. 3. palaeogeographic configuration of the north sea basin and the surrounding seas during the middle to late paleocene. from ziegler (1988, 1990). 17 the paleocene do not entirely concur with the conclusions of berggren & aubert (1975, pp. 73, 182) based on the palaeobiogeographic analysis of the paleocene benthonic foraminifera faunas. they claimed that the “geographic distribution of most of the elements of the (paleocene benthonic foraminifera) assemblages was essentially cosmopolitan” and was “attributed to more equitable climatic conditions (lower polar – equatorial thermal gradient) and warmer, more uniform thermal structure of the oceans” and that “this distribution is apparently independent of taxonomy”. they recognised two basic faunal assemblages attributed to palaeobathymetry, i.e. the shallow-water midway and the deep-water velasco types. a preliminary interpretation of the otolith data as outlined above suggests that climatic differences played a bigger role in the composition of fish assemblages. paleocene/eocene boundary during early eocene the complete (or partial?) separation of the north sea basin from the warmer atlantic seas in the south-west disappeared (figs 3, 6). as the western approaches basin (off normandy, france) became connected with the southern north sea basin (bonde 1979; ziegler 1982), water current circulation changed and warm subtropical to tropical water was able to flow in from the west. this is well expressed in buchardt’s (1977, 1978) palaeotemperature curve and it also brings along a sudden change in the composition of the fish fauna as documented by otoliths. from thanetian through ypresian to lutetian gadiform otoliths and argentinidae disappear almost entirely, whereas ophidiiform and perciform otoliths became extremely common and rich in species. the faunal composition of fish changed from warm temperate to tropical. ★ ★ ★ d c m fig. 4. palaeogeography and palaeoecology in the southern north sea basin during the campanian to danian. from ziegler (1982, 1988). open shelf carbonate environment (brickwork pattern) prevailed over much of north-west europe; otolith localities are marked with an asterisk. c: campanian otolith locality (erratic boulders, voigt 1926), d: danian (from fakse, this bulletin). m: maastrichtian (bavaria; unpublished data, w. schwarzhans). 18 mo-clay (upper thanetian to lower ypresian) there is one important osteologically based teleost fauna to be discussed in this context, i.e. the wellknown mo-clay deposit from northern denmark (jutland) which is transitional in age from the paleocene to the eocene (fig. 5). stratigraphically and genetically, this is probably the most closely referable osteologically based fauna from the north sea basin for correlation with the selandian otolith associations of copenhagen. unfortunately, the aragonitic otoliths are dissolved in these diatomitic/volcanic ash sediments so that skeletons cannot be directly correlated with otoliths in situ. the mo-clay fishes have never been comprehensively described, but bonde (1966) has published an extensive species list with some later amendments (bonde 1979). so far, the fauna list comprises 29 species of teleost fishes. the two most common species in this list are an argentinoidei and an osmeroidei. other common moclay fishes reported are an aulostomoidei, a polymixiidae, a percoidei and a small scorpaenidae, which bonde regards as possibly belonging to the fossil genus ampheristus (now conclusively placed in ophidiids, both by osteology and otoliths) described from the london clay. in this respect it must be noted that the abundance of ophidiid otoliths in palaeogene sediments represents one of the most serious discrepancies to theosteological findings. in fact, ophidiid fish skeletons have hardly been reported at all. other fishes listed by bonde include gadiformes (one merlucciidae and one gadidae), an elopidae (or pterothrissidae?), a clupeidae, an osteoglossiformes, an anguilliformes, a veliferidae, ★t s s t ★ ★ ★ s+t★ t fig. 5. palaeogeography and palaeoecology in the southern north sea basin during the selandian and thanetian: closed basin stage; dots mark near shore sandy facies, hatched areas are lower shelf clayey environment and horizontal lines are open marine environment; otolith localities are marked with an asterisk. s: selandian; copenhagen (koken 1885) and this study; erratic boulders from north-east germany (roedel 1930); undescribed material from bavaria. t: thanetian; england (stinton 1965, 1977); belgium (nolf 1978); undescribed material from bavaria; the small fish marks the mo-clay area with skeleton findings (bonde 1966, 1979). 19 a lamproidei, a zeidae, and three percoidei including a carangidae, several scombroidei, a centrolophidae and a nomeidae. these osteological findings from the mo-clay correlate quite well with the otolith findings of the selandian and the common argentinoid and osmeroid skeletons of the mo-clay may compare to the argentina and protargentinolithus otoliths of the selandian. the ampheristus skeletons could relate to the ophidiid otoliths. skeletons and otoliths both represent several primitive percoids, for example of the family carangidae. merlucciidae and gadidae are known both from skeletons and otoliths. however, otolith findings from the selandian are much more common and more diverse, including three records of the family macrouridae, which is not represented in any of the other european palaeogene faunas, be it otolith or skeleton based. other matches could be in the following groups: elopidae / pterothrissidae, clupeidae, anguilliformes, veliferidae, zeidae and centrolophidae / nomeidae. osteological findings with lacking otolith matches are the aulostomoidei (otoliths of fishes of this group are extremely small and so far have not been recorded as fossils), scombroidei (generally rare as otoliths in the fossil record) and the polymixiidae. likewise there are also some common groups represented by otoliths, which do not seem to find their counterparts in the skeleton record, i.e. ariidae, chlorophthalmidae, macrouridae and berycidae (hoplostethus and centroberyx). in general, the mo-clay fish fauna has more in common with the selandian–thanetian otolith association ★ ★ ★ ★ ★ fig. 6. palaeogeography and palaeoecology in the southern north sea basin during the early to middle eocene: southern north sea reopened to the atlantic ocean through a ‘channel’; warm water influx characterised by limestone / sandy limestone formations in the south-west; elsewhere legend as in figs 4, 5; asterisk marks otolith localities. 20 than with the tropical eocene otolith assemblages. despite the observation that ‘skeleton’ and ‘otolith’ findings largely complement each other (nolf 1985) it is found that in this case they correlate surprisingly well (with some exceptions, see above). natural causes that affect the discrepancy between otolith and skeleton findings include stratigraphy and facies but in addition it is likely that some of the differences are caused by how skeletons and otoliths are being identified systematically. in conclusion, it seems that the major change in the fish fauna of the north sea basin has occurred in post mo-clay times. this indicates that the warm-water influx through the opening of the western approaches basin (off normandy, france) and the ‘strait of dover’ connecting the north sea basin with the north-eastern part of the atlantic ocean either was established later or at least at the time of the terminal paleocene was not fully effective. evolutionary interpretation the turn from the cretaceous to the tertiary times has long been recognised as important in the evolution of the teleostei and mainly at the suprageneric level (suborders, families and sometimes genera). otolith data from this time interval are growing. these data are from the late cretaceous of north america (nolf & dockery 1990; nolf & stringer 1996), bavaria in southern germany (unpublished data, w. schwarzhans), from the paleocene of north america (nolf & dockery 1993), denmark (this bulletin) and west greenland and bavaria (unpublished data, w. schwarzhans). it is therefore not surprising that the paleocene of denmark yielded a number of first records of otoliths such as for the anguillidae, clupeidae, salmonidae, ?percopsiformes, scorpaenidae, gempylidae, centrolophidae, ostraciidae and many gadiformes. the ostraciid ostracion pergravis represents the first ever otolith based record of the order tetraodontiformes. since tetraodontiformes are assumed to have derived from perciformes and in particular from the advanced perciform suborder acanthuroidei, this early and morphologically clear-cut otolith finding gives some doubt to the true origin of the tetraodontiformes. in this respect it must be noted that patterson (1964, p. 470), based on osteological findings, discussed the relationship of acanthuroidei with the extinct beryciform family pharmacichthyidae from the late cretaceous. further very surprising evidence from the paleocene of denmark is the abundance and richness of the gadiform fishes as they are represented by the otoliths. the families ranicipitidae, lotidae, gadidae and macrouridae are here represented with their earliest records; the latter with early representatives of three modern lineages (the genera coelorhynchus, hymenocephalus and coryphaenoides). cretaceous gadiforms so far have not been recorded, neither by osteological nor by otolith findings (in fact, the only cretaceous records of the paracanthopterygii are based on otoliths from the ophidiiformes; unpublished data, w. schwarzhans). the wide range of gadiform morphologies in the paleocene otolith findings indicate that this order must have been of pre-tertiary origin and is expected to reach much further backwards in the geological history than previously reported. in the summary on the maastrichtian otoliths from bavaria schwarzhans (1996) presented a scheme composed of four morphologic-evolutionary categories. the four categories are: category 1: persistent taxa category 1 includes otoliths with morphologies that have not altered significantly since the paleocene times. these otoliths can reasonably well be assigned to living genera (or sometimes listed with a sensu lato). the living genera that are represented by otoliths from the paleocene in denmark are conger, rhechias (congridae), arius (ariidae), argentina (argentinidae), aulopus (aulopidae), chlorophthalmus (chlorophthalmidae), raniceps (ranicipitidae), molva (lotidae), coelorhynchus, hymenocephalus, coryphaenoides (macrouridae), bidenichthys (bythitidae), hoplostethus, centroberyx (berycidae), scorpaena (scorpaenidae), acro21 poma (acropomatidae), mupus (centrolophidae) and ostracion (ostraciidae). the total is 20 species out of 44 or 45% representing persistent genera. in comparison only 18% of such species have been identified in the maastrichtian of bavaria indicating that a rather drastic raise of reasonably generically attributable otolith morphologies occurred across the cretaceous–palaeogene boundary. the increase in category 1 from maastrichtian to paleocene is mainly due to additions not replacements, since most of the maastrichtian category 1 forms are still present in the paleocene of denmark (for instance congridae, argentinidae, chlorophthalmidae, bythitidae, berycidae). category 2: extinct early specialised taxa category 2 contains highly specialised morphologies without apparent affinities to living taxa. such otoliths are interpreted to represent extinct taxa that are more or less indigenous to the late cretaceous (and sometimes to the paleocene as well). their systematic placement is often problematical. representatives in the paleocene of denmark are few: genartina (near pterothrissidae), genus myctophidarum schnetleri (myctophidae), genus bythitidarum rasmussenae (bythitidae) together with three forms left in open nomenclature. the total is six species or 14%, which compares to 32% from the maastrichtian of bavaria. most remarkable is the total lack of extinct specialised beryciformes that form the majority of morphologies in this category in the late cretaceous, both in osteological and otolith findings. category 3: extinct plesiomorphic taxa category 3 includes otoliths with plesiomorphic morphologies that are usually attributable to extant families, but sometimes can be of problematical allocation when very generalised in appearance. representatives for category 3 in the paleocene of denmark are pteralbula (pterothrissidae), protargentinolithus (argentinidae), palaeogadus (merlucciidae), gadomorpholithus (lotidae), protocolliolus (gadidae), genus ophidiidarum seelandicus, genus veliferidarum harderi and genus stromateoidarum sp. the total is nine species or 21%. compared to the maastrichtian of bavaria, which comprises 41%, this again documents a decrease of forms within this category across the cretaceous–palaeogene boundary. it is also a change in quality; the wide range of maastrichtian morphologies in this category is merely represented by three remaining forms in the paleocene, namely pteralbula, an ophidiid and a possible veliferid. the other paleocene forms within category 3 are all newcomers. category 4: ‘missing links’ under this informal heading i had grouped certain plesiomorphic otolith morphologies, which are believed to be situated near major dichotomical events (suborder to family level) in the phylogeny of persistent, living teleosts. this category may seem weakly defined and in fact the distinction into categories 3 and 4 is quite fluent and subject to changes with the increasing status of knowledge on the one hand and interpretative alterations on the other. i have tentatively placed in this category the following species from the paleocene of denmark: genus anguillidarum semisphaeroides, genus anguillidarum sp., genus clupeidarum rectiventralis, genus salmonidarum acutirostratus, genus ?percopsiformorum enigmaticus, genus zeiformorum janni, genus carangidarum sp., genus sparidarum sp. and genus gempylidarum merus. the total is nine species or 21% and the corresponding number from the maastrichtian of bavaria, germany is 9%. typically for the category 4 is that most species have to be left in open generic nomenclature because of the generalised morphology that often looks like an ‘archetypical’ catch-all for the family or order/ suborder in question. many of these records are also the earliest otolith findings for the respective families. finally, the large amount of perciformes in this category is remarkable and probably has to do with the replacing of the late cretaceous beryciform dominated teleost fauna with primitive perciforms during early tertiary. in conclusion, it can be stated that a major evolutionary change in the composition of the teleost fauna has occurred across the cretaceous–palaeogene boundary as evidenced by otolith findings. this is mainly due to the change from a late cretaceous beryciform dominated teleost fauna to a paleocene gadiform and perciform dominated association. since many of the late cretaceous beryciforms have been categorised as early specialised extinct taxa, this change brings along a sudden decrease of otolith morphologies of the evolutionary category 2. other evolutionary categories – 22 persistent taxa (category 1), extinct plesiomorphic taxa (category 3) and ‘missing links’ (category 4) – increase comparatively. this reflects the introduction of primitive early morphologies (categories 3 and 4) in many families and at the same time the increase of morphologies that can be reasonably attributed to living genera (category 1). stratigraphic significance time interval. common selandian otolith species that so far have definitely not been reported from the thanetian include rhechias angulosus, protargentinolithus balticus, p. procerus, chlorophthalmus postangulatus, gadomorpholithus ponderosus, protocolliolus amorphus, coelorhynchus balticus and genus veliferidarum harderi. of greater interest at this stage are the species of the less known thanetian otoliths that are not present in theselandian.these are hoplobrotula protensa, holocentrus sheppeyensis and genus serranidarum serranoides. the use of the otolith findings from the paleocene of denmark for stratigraphic purposes is still very limited with the data available. for the time being the large selandian fauna stands somewhat isolated from otolith findings in adjacent stratigraphic units. the number of species that the selandian otolith assemblage has in common with the smaller otolith associations of danian and thanetian age is already impressive and as more otoliths will be described from these two respective strata this number is likely to increase. only few selandian species may remain indigenous to this systematic taxonomy a review of the type material relevant to this study is presented below. koken (1885, 1891). koken’s type material is from the selandian of copenhagen. the material is currently under revision by the present author and d. nolf. koken’s figures and descriptions are very accurate and the species established by him are all valid. however, some of his generic allocations need revision, due to the lack of extant comparative otoliths available at the time of his study. this revision is presented in the following summary and given in table 1. roedel (1930). roedel (1930) is the second author who previously described otoliths from the selandian of the north sea basin. roedel (1930) obtained his otoliths from erratic boulders in north-east germany. andersen & heilmann-clausen (1984) showed that these erratic boulders are equivalent to the lellinge greensand in denmark. roedel (1930) did not reach the same high standard as was previously set by koken (1885, 1891) and it has been known since a long time that roedel’s identifications needed a thorough revision. until recently, his collection was considered to be lost during the second world war, but recently the complete typecollection was discovered in the collection of the humboldt-university, berlin by w.-d. heinrich. the material has been incorporated in this study. many of roedel’s identifications were based on inadequate material, i.e. eroded and non-diagnostic juvenile or fragmented specimens. because the drawings published by roedel are schematic, the type specimens are redrawn in this bulletin whenever it was found appropriate.the results of the revision of roedel’s original material are summarised in table 2. other material. few additional publications on otoliths from other strata or areas have an impact on the 23 koken (1891) this study arius danicus koken 1891 arius danicus koken 1891 genus gadidarum ponderosus koken 1885 gadomorpholithus n. gen. ponderosus (koken 1885) merluccius balticus koken 1885 coelorhynchus balticus (koken 1885) genus apogonidarum lacinatus koken 1885 hoplosthetus lacinatus (koken 1885) genus apogonidarum integer koken 1885 centroberyx integer (koken 1885) trachinus seelandicus koken 1885 genus ophiidarum seelandicus (koken 1885) genus inc. sed. conchaeformis koken 1885 pteralbula conchaeformis (koken 1885) table 1. koken’s original names and the revised names used in this study roedel’s (1930) identification this study arius germanicus koken 1891; synonym of arius danicus koken 1891 non koken 1891 arius rotundus roedel 1930 synonym of arius danicus koken 1891 merluccius schmitti roedel 1930 synonym of coelorhynchus balticus (koken 1885) merluccius globulosus roedel 1930 juvenile otolith, likely synonym of coelorhynchus balticus (koken 1885) merluccius latisculptus roedel 1930 a fragmented ophidioid otolith, most likely synonym is genus ophidiidarum seelandicus (koken 1885) merluccius nanus roedel 1930 a juvenile and fragmented gadiform otolith without real diagnostic features; it is recommended not to use this species name genus gadidarum insuetus roedel 1930 a mollusc fragment, not an otolith solea solitarius roedel 1930 a strongly eroded, non-identifiable juvenile otolith; certainly not a pleuronectiform; it is recommended not to use this species name genus berycidarum marchicus roedel 1930 genus scorpaenidarum marchicus (roedel 1930). this species is not known from the danish localities genus berycidarum balticus roedel 1930 protargentinolithus n. gen. balticus (roedel 1930) genus percidarum holsaticus roedel 1930 the origin of this otolith is questionable; it closely resembles small specimens of brachydeuterus gaemersi, which is well known from upper oligocene erratic boulders from the same area (‘sternberger gestein’); it is recommended not to use this name genus percidarum erraticus roedel 1930 argentina erratica (roedel 1930); an annotated paratype belongs to protargentinolithus balticus. however, this annotation may well be in error, since roedel reported only one single specimen genus percidarum obliquestriatus roedel 1930 genus anguilliformorum obliquestriatus (roedel 1930). this specimen exhibits a plesiomorphic otolith pattern, which is very difficult to assign to either a genus or family. this species is not known from the danish localities genus percidarum minimus roedel 1930 an eroded, juvenile perciform otolith without real diagnostic characteristics; it is recommended not to use this species name genus sparidarum gregarius koken 1891; a fragmented specimen, probably a synonym of protoargentinolithus balticus non koken 1891 (roedel 1930) genus inc. sed. erhardvoigti roedel 1930 an eroded specimen, but apparently a synonym of pteralbula conchaeformis (koken 1885) table 2. roedel’s species names and the names used in this study 24 stage and locality danian selandian species fakse gemmas allé sundkrogen vestre gasværk kongedyb total pterothrissidae pteralbula conchaeformis (koken 1885) 4 165 306 10 485 genartina hauniensis n. sp. 94 3 97 anguillidae genus anguillidarum semisphaeroides n. sp. 3 3 genus anguillidarum sp. 1 1 congridae conger illaesus n. sp. 1 1 2 rhechias angulosus n. sp. 1 1 12 23 3 40 clupeidae genus clupeidarum rectiventralis n. sp. 4 4 genus clupeidarum aff. rectiventralis n. sp. 2 2 ariidae arius danicus koken 1891 4 12 50 1 67 salmonidae genus salmonidarum acutirostratus n. sp. 1 1 argentinidae protargentinolithus balticus (roedel 1930) 351 252 21 624 protargentinolithus procerus n. sp. 96 64 6 166 argentina erratica (roedel 1930) 1 176 20 10 207 argentina longirostris n. sp. 62 4 66 genus argentinidarum sp. 6 1 7 aulopidae aulopus tortus n. sp. 4 4 chlorophthalmidae chlorophthalmus postangulatus n. & d. 1993 3 4 46 113 3 169 myctophidae genus myctophidarum schnetleri n. sp. 15 2 17 genus myctophidarum sp. 2 1 3 percopsiformes genus ?percopsiformorum enigmaticus n. sp. 5 1 6 ranicipitidae raniceps hermani nolf 1978 3 3 merlucciidae palaeogadus sinangulatus n. sp. 6 26 116 11 159 lotidae gadomorpholithus ponderosus (koken 1885) 4 19 63 1 87 molva palaeomorpha n. sp. 4 4 8 gadidae protocolliolus amorphus n. sp. 7 3 98 108 macrouridae coelorhynchus balticus (koken 1885) 7 87 371 9 474 hymenocephalus rosenkrantzi n. sp. 9 7 16 coryphaenoides amager n. sp. 2 8 10 gadiformes spp. 4 7 447 246 30 734 ophidiidae genus ophidiidarum seelandicus (koken 1885) 18 130 538 10 696 bythitidae bidenichthys lapierrei (nolf 1978) 71 2 2 2 77 genus bythitidarum rasmussenae n. sp. 5 5 veliferidae genus veliferidarum harderi n. sp. 1 37 42 1 81 zeiformes indet. genus zeiformorum janni n. sp. 1 1 berycidae hoplostethus lacinatus koken 1885 1 35 64 3 103 centroberyx integer (koken 1885) 2 26 27 55 centroberyx fragilis n. sp. 89 3 55 86 4 237 table 3. fish otolith species and number of specimens from the paleocene of denmark 25 identification of the danish paleocene otoliths. these include stinton (1965), who described otoliths from the thanetian of the london basin, nolf (1978), who reported on otoliths from the thanetian of belgium and nolf & dockery (1993), who described paleocene otoliths from alabama, usa. the latter authors referred the alabama collection to the danian stage, but it may well belong to the selandian stage (i.e. biozone np4; nolf, personal communication 1999). the material from england, belgium and the usa is not reviewed here, but reference is made wherever it is appropriate. unpublished collections of paleocene otoliths from west greenland are also referred to here, but is not dealt with further in this bulletin. at this stage it should be noted that the greenland fauna is remarkably similar in composition and possibly also at the species level to the danish selandian otoliths. a further extensive paleoceneotolithcollection frombavaria, southern germany and presently under investigation by schwarzhans is remarkable for its difference to the danish otolith assemblage. repository all the described and illustrated otoliths (table 3) including holotypes and paratypes are stored and registered at the geological museum of the university of copenhagen, denmark and are indicated with the prefix mguh. other specimens that originated from harder’s and rosenkrantz’ collections are also kept at the geological museum of the university of copenhagen but without mguh numbers. additional specimens inspected here are from m. nielsen, mrs. a. rasmussen and k. schnetler’s collections. the specimens are respectively all kept in schnetler’s and mrs. a. rasmussen private collection in denmark. koken’s (1885) type material from the paleocene of copenhagen is also kept at the geological museum of the university of copenhagen. roedel’s (1930) type material from the paleocene erratic boulders of cöthen, north-east germany are in the collection of the palaeontological museum of the humboldt-university in berlin (pmhub), but do not have reference numbers. taxonomic description in the taxonomic description the morphological terminology follows that of koken (1884) with the amendments proposed by weiler (1942) and schwarzhans (1978). open generic nomenclature is used for species of uncertain generic position and follows the recommendations made by nolf (1985). new fossil otolith based generic names are introduced and the formal ending -lithus is used to indicate that the new stage and locality danian selandian species fakse gemmas allé sundkrogen vestre gasværk kongedyb total scorpaenidae scorpaena corallophilus n. sp. 35 35 apogonidae genus apogonidarum sp. 2 2 acropomatidae acropoma sp. 11 2 6 19 carangidae genus carangidarum sp. 62 1 63 sparidae genus sparidarum sp. 2 5 2 9 gempylidae genus gempylidarum merus n. sp. 3 3 centrolophidae mupus sinuosus (stinton 1965) 45 3 48 genus stromateoidarum sp. 7 7 ostraciidae ostracion pergravis n. sp. 7 7 number of species 13 13 34 31 21 44 number of specimens 222 68 2062 2528 140 5020 table 3 (continued) 26 fossil genus is based on otoliths. this procedure is used toavoidconfusionwith skeleton-based fossil teleost genera, which could occur when cited separately. it is assumed that the two recent sciaenid genera otolithes and pseudotolithus are common knowledge. i strongly recommend using this taxonomic ‘marker’ in future descriptions of new fossil otolith-based genera, when they are established. new fossil otolith based generic names are here introduced in two instances of species apparently representing fossil genera, i.e. protargentinolithus and gadomorpholithus. the classification used here follows the one proposed by nelson (1994). explanation of abbreviations used under the headings measurements: length = l, height = h, thickness = t, length:height = l:h, height:thickness = h:t, affinity = aff., not measured = n.m., paratype = para., holotype = holo. all measurements in the tables are given in millimetres. order elopiformes greenwood et al. 1966 family pterothrissidae gill 1893 remarks. pterothrissid otoliths form a common and widespread faunal component in the late cretaceous and the early tertiary fish faunas, although mostly represented by just one or two species at each locality. in the recent, the two existing species are endemic and are restricted to the continental slopes of west africa (pterothrissus belloci) and japan (pterothrissus gissu). genus pteralbula schwarzhans 1981b type species. genus inc. sed. conchaeformis koken 1885. pteralbula conchaeformis (koken 1885) fig. 7a–j 1885 genus inc. sed. conchaeformis koken, p. 113, plate 5, fig. 25. 1930 genus inc. sed. erhardvoigti n. sp. roedel, p. 67, plate 1, fig. 14. material. 485 otoliths from the selandian, paleocene: 4 small specimens from gemmas allé, 165 specimens from sundkrogen (mostly small, 10 large specimens), 306 specimens from vestre gasværk (including 148 large specimens), 10 small specimens from kongedyb; figured specimens from sundkrogen (mguh 26037– 26042). in addition, roedel’s holotype of genus inc. sed. erhardvoigti has been inspected and is redrawn in fig. 7c. measurements l h t l:h h:t 8.75 6.25 2.20 1.40 2.80 6.75 5.00 1.75 1.35 2.85 6.25 3.80 1.30 1.65 2.90 holo.* 4.50 2.90 1.05 1.55 2.75 2.95 1.85 n.m. 1.60 1.85 1.05 n.m. 1.75 for abbreviations used in the table, see opposite. * of p. erhardtvoigti. description. the otoliths are large and up to 10 mm or more in size. they are rounded to rectangular in outline and with a pronounced postdorsal angle in the adults. the adult otoliths have an inner face that is rather smooth and markedly convex, which is characteristic for the otoliths of the fossil otolith-based genus pteralbula. the sulcus is oriented diagonally on the inner face and with a tapering narrow and rather long cauda and a wider ostium, which is reaching close to but is not opened towards the antero-dorsal rim. remarks. juvenile otoliths that are less than 5 mm in length prevail, which is common for the occurrence of the fossil pterothrissid otoliths. the juveniles show a much stronger ornamentation on the inner and outer faces and have a more regularly curved outline than the adults. like in many fossil pterothrissids their otolith morphology is very generalised without any real diagnostic features. such specimens can only be reliably identified in the presence of a suitable assemblage of otoliths of different ontogenetic stages (schwarzhans 1981b). family indet. genus genartina frizzell & dante 1965 type species. genus inc. sed. hampshirensis schubert 1916. genartina hauniensis n. sp. fig. 8a–h 27 a b c d e g i j h f fig. 7. pteralbula conchaeformis (koken 1885). a, b, d–j: selandian, sundkrogen, mguh 26037–26042, × 8. c: refigured holotype of inc. sed. erhardvoigti roedel 1930 (pmhub), selandian, erratic boulders in northern germany, × 12. 28 type locality. sundkrogen, excavation 1920. type stratum. lellinge greensand, selandian, paleocene. derivation of name. after hafnia, the latin name for the city of copenhagen, from where the type material is obtained. holotype. fig. 8b, c, mguh 26043. paratypes. fig. 8a, d–h, mguh 26044–26048. diagnosis. high bodied, thin flat otoliths with a gentle and deeply curved ventral rim and a short, sharp rostrum. the dorsal rim is expanded with a characteristic deep and angular antero-dorsal incision. further material. 88 specimens from sundkrogen, 3 specimens from vestre gasværk. measurements l h t l:h h:t ~ 3.50 3.10 n.m. para. ~ 3.30 2.85 0.50 ~ 1.10 5.7 holo. ~ 2.80 2.60 n.m. para. ~ 2.30 2.05 n.m. para. 1.80 1.60 0.30 1.10 5.3 para. 1.45 1.30 n.m. 1.10 para. for abbreviations used in the table, see p. 26. description. otoliths are high bodied, thin and up to 4 mm in size. the length:height index is about 1. the ventral rim is smooth, deep and gently curved. the dorsal rim is strongly expanded medially and with a a b c d e f gh fig. 8. genartina hauniensis n. sp. a, d–h: paratypes, selandian, sundkrogen, mguh 26044–26048, × 20. b, c: holotype, selandian, sundkrogen, mguh 26043, × 20. 29 conspicuous, deep and angular anterior and a more gentle posterior incision. the anterior rim has a short but sharp and very fragile rostrum. there is no excisura or antirostrum. the posterior rim is rounded, smooth and somewhat projecting behind the caudal tip. the inner face is almost flat with a slightly supramedian positioned sulcus. the cauda is tapering, long and somewhat deepened, reaching very close to the posterior tip of the otolith. the ostium is slightly widened and somewhat more deepened, opening anteriorly below the antero-dorsal incision. the dorsal depression is very large and occupies almost the entire expanded area of the dorsal field. the ventral line is feeble or absent. the outer face is smooth and slightly convex. all rims are sharp. ontogeny and variability. small otoliths, i.e. of 1.5 mm and less in size exhibit a fine marginal ornamentation on the inner and outer faces. the postdorsal concavity may not be developed and also the predorsal incision may be less intense. the posterior rim can be rather blunt. in terms of variability otoliths of the same size are morphologically very similar. discussion. g. hauniensis resembles the younger g. hampshirensis, which is known under various names (see nolf 1985 for synonymy), from the thanetian to late eocene of england and belgium and the two species are probably closely related. the otoliths that are referred to g. hauniensis however are easily recognised by their very specific development of the dorsal rim and the very small length:height index. large otoliths are also characterised by the absence of any marginal ornamentation. apart from the two species mentioned above g. texana from the eocene of the usa is the only additional species of the genus. the zoological relationship of the fossil otolith-based genus genartina is obscure. it has often been associated with argentinidae or osmeridae (stinton and nolf, various publications). here, it is tentatively interpreted as an aberrant representative of an extinct family of the elopiformes. order anguilliformes regan 1909 suborder anguilloidei regan 1909 family anguillidae rafinesque 1810 genus indet. genus anguillidarum semisphaeroides n. sp. fig. 9a–d type locality. vestre gasværk. type stratum. selandian, paleocene. derivation of name. from semisphaeroides (greek) = hemispherical; referring to the thick appearance of the otolith, i.e. with a strongly convex inner face and a nearly flat outer face. holotype. fig. 9a–c, mguh 26049. paratype. fig. 9d, mguh 26050. diagnosis. thick and rounded otoliths with a strong convex inner face and a flat outer face. the sulcus is long and deep with a long cauda and a short ostium, which is open to the anterior. there is no dorsal depression or ventral furrow. further material. vestre gasværk: 1 specimen. measurements l h t l:h h:t 3.65 2.80 1.45 1.30 2.1 holo. 2.35 2.00 n.m. 1.15 para. for abbreviations used in the table, see p. 26. description. the otoliths are thick and with a rounded outline; they are massive and with sizes up to about 4 mm. all rims are smooth, gently curved and without prominent angles. the dorsal rim has a strongly rounded postdorsal angle. the inner face is strongly convex and smooth; it is without a dorsal depression or ventral furrow except for the deep sulcus. the sulcus has a long, narrow and very deep cauda and a short, dorsally widened and shallower ostium, which is opening towards the antero-dorsal rim. the ostial channel is indicated. the colliculi are fused and not reduced in length or width. the outer face is flat and smooth. otolith rims are thick. variability. the holotype is slightly more elongated than the two paratypes. discussion. the illustrated specimens are in a perfect condition. the characters given in the diagnosis easily identify these otoliths. the deep and anteriorly open 30 sulcus resembles otoliths of the family anguillidae and certain genera of the family ophichthyidae (for example echelus). it is likely that genus anguillidarum semisphaeroides n. sp. represents an extinct genus (or family) within or near to the anguillidae. two more species fromtheeocene of belgium and england are known in the fossil record and are somewhat similar in habitus and might indeed be related. one is anguilla rectangularis stinton & nolf 1969, which is characterised by its rectangular outline. the other one is echelus contractus stinton 1975, which is a species characterised by otoliths that are more elongated and with a flat dorsal rim and a more pronounced postdorsal angle. genus anguillidarum sp. fig. 9e–g material. vestre gasværk: 1 specimen, selandian, mguh 26051. measurements l h t l:h h:t 2.55 1.55 0.55 1.65 2.8 for abbreviations used in the table, see p. 26. description. the small otolith is similar to genus anguillidarum semisphaeroides. it is however characterised by being thinner, having an elongated appearance and it has a rather flat ventral rim. discussion. this single specimen probably represents a distinct species that is closely related to genus anguillidarum semisphaeroides n. sp. suborder congroidei regan 1909 family congridae kaup 1856 genus conger oken 1817 a b c d e f g fig. 9. a–d. genus anguillidarum semisphaeroides n. sp. a–c: holotype, selandian, vestre gasværk, mguh 26049; a: × 20; b, c: × 12. d: paratype, selandian, vestre gasværk, mguh 26050, × 20. e–g. genus anguillidarum sp. selandian, vestre gasværk, mguh 26051, × 20. 31 a b c d f g h i j k l m n e fig. 10. a–e. conger illaesus n. sp. a–c: holotype, selandian, vestre gasværk, mguh 26052, × 12. d, e: paratype, selandian, sundkrogen, mguh 26053, × 12. f–l. rhechias angulosus n. sp. f–h: holotype, selandian, vestre gasværk, mguh 26054, × 12. i, j: paratype, selandian, sundkrogen, mguh 26055, × 20. k: paratype, danian, fakse, mguh 26057, × 20. l: paratype, selandian, gemmas allé, mguh 26056, × 20. m, n. genus anguilliformorum obliquestriatus (roedel 1930). refigured holotype from erratic boulders of northern germany (pmhub), selandian, × 20. 32 type species. muraena conger linnaeus 1758. conger illaesus n. sp. fig. 10a–e type locality. vestre gasværk. type stratum. selandian, paleocene. derivation of name. illaesus (latin) = unharmed, intact. holotype. fig. 10a–c, mguh 26052. paratype. fig. 10d, e, from sundkrogen, mguh 26053. diagnosis. elongated, nearly symmetrical and massive otoliths. the sulcus is undifferentiated and filled with a single, oval colliculum, which is somewhat reduced to the anterior. the dorsal depression is wide and prominent. measurements l h t l:h h:t 3.10 1.60 0.85 1.95 1.9 holo. 2.75 1.40 0.55 1.95 2.5 para. for abbreviations used in the table, see p. 26. description. otoliths are rather elongated, thick, massive and almost symmetrical in outline. the ventral rim is smooth and shallow. the dorsal rim is shallow, somewhat undulated and occasionally it has a somewhat pronounced postdorsal angle. the anterior and posterior tips are slightly pointed; the anterior tip is sharper than the posterior tip. the inner face is rather flat with a slightly inclined wide and shallow sulcus, which is anteriorly reduced, but reaches close to the antero-dorsal rim. the ostium and cauda is not differentiated; the sulcus is filled with a large uniform colliculum, which is slightly reduced in length at its anterior and posterior tips. the ostial channel is short and distinct. the dorsal depression is wide and pronounced. there is no ventral furrow. the outer face is more strongly convex than the inner face and smooth. the rims are moderately sharp and smooth. variability. the paratype is slightly smaller and thinner than the holotype. also it shows a clear postdorsal angle, which indicates a certain level of variability. discussion. this is a typical otolith of the genus conger. otoliths referred to conger illaesus are easily recognised by the combination of characters given in the diagnosis. there are no comparable species recorded from the paleogene of europe. genus rhechias jordan 1922 type species. rhechias armiger jordan 1922. rhechias angulosus n. sp. fig. 10f–l type locality. vestre gasværk. type stratum. selandian, paleocene. derivation of name. angulosus (latin) = angular; referring to the prominent postdorsal angle. holotype. fig. 10f–h, mguh 26054. paratypes. fig. 10i, j, from sundkrogen, mguh 26055; fig. 10k, from fakse, mguh 26057; fig. 10l, from gemmas allé, mguh 26056. diagnosis. moderately elongated and thick otoliths with pointed anterior and posterior tips and a prominent and sharp postdorsal angle. the sulcus is shallow, narrow and rather short and inclined at about 5 to 10°. the ostium is strongly reduced anteriorly and fused colliculum terminating at some distance from the anterior rim. the dorsal depression is wide and marked. further material. vestre gasværk: 22 specimens; sundkrogen: 12 specimens and kongedyb: 3 specimens. measurements l h t l:h h:t 5.40 3.30 1.40 1.65 2.35 holo. 4.55 2.75 n.m. 1.65 para. 2.90 1.85 0.75 1.55 2.45 para. 2.75 1.75 n.m. 1.60 para.* 1.95 1.25 n.m. 1.55 para. for abbreviations used in the table, see p. 26. * fakse. description. the otoliths are moderately elongated with a length:height index of 1.55 to 1.60. the size is up to about 5 mm. otoliths are thick with a height:thickness 33 index of about 2.5. the ventral rim is gently curved and deepest to the anterior of the mid-part of the otolith. the dorsal rim has a faint predorsal and a very prominent and sharp postdorsal angle. anterior and posterior tips are more or less pointed. all rims are smooth. the inner face is convex with a rather shallow, short and narrow sulcus inclined at 5 to 10°. the cauda is straight and with a rounded tip that terminates at some distance from the posterior rim. the ostium is reduced; it is not open to the anterior but instead terminates at some distance from the anterior rim. the ostial channel is absent or extremely feeble and reduced. the colliculum is fused, short and reduced to the anterior. the dorsal depression is wide, marked and not very deep. the ventral line is absent or very feeble close to the ventral rim. the outer face is convex and smooth. ontogeny and variability. the smallest specimen is about 2 mm in length and resembles the larger ones except for having somewhat more rounded anterior and posterior tips. likewise, variability seems to be restricted to small differences in the outline and the expression of the anterior part of the sulcus. discussion. the anteriorly reduced and straight sulcus in combination with the marked dorsal depression is regarded as characteristic for the species of this genus. congrid otoliths are well known from the early tertiary of europe and elsewhere. rhechias angulosus n. sp. is well distinguished by the characters given in the diagnosis from rhynchoconger eocenicus from the early eocene of england (shepherd 1916) and rhynchoconger sp. from the paleocene of alabama, usa (nolf & dockery 1993). figure 10m, n depicts the unique holotype of the genus anguilliformorum obliquestriatus (roedel 1930) from the time equivalent erratic boulders in north-east germany for comparison. this species apparently is unrelated to any of the anguilliform species described above from the paleocene of denmark. order clupeiformes bleeker 1859 family clupeidae cuvier 1817 genus indet. genus clupeidarum rectiventralis n. sp. fig. 11a–i type locality. sundkrogen. type stratum. selandian, paleocene. derivation of name. rectiventralis (latin), referring to the straight ventral rim of the otolith. holotype. fig. 11a–c, mguh 26058. paratype. fig. 11d, e, mguh 26059. diagnosis. the otolith is elongated, thin and fragile. the ventral rim is straight. the rostrum is prominent. the sulcus is wide, long, deep and with a cauda that is almost as long as the ostium. further material. 2 specimens; fig. 11f–i, 2 juvenile specimens (listed as genus clupeidarum aff. rectiventralis in table 3), mguh 26060–26061. measurements l h t l:h h:t 1.60 0.85 0.20 1.95 4.2 holo. 1.30 0.80 0.25 1.60 3.2 para. 1.30 0.85 0.25 1.55 3.2 (aff.) 1.00 0.70 0.20 1.45 3.5 (aff.) for abbreviations used in the table, see p. 26. description. the otoliths are quite small, thin, fragile and elongated with a length:height index of almost 2. the ventral rim is straight, horizontal and somewhat undulating. the dorsal rim is also nearly flat, smooth and without prominent angles. the posterior rim is gently rounded with a faint incision just above the middle of the specimen. the anterior rim depicts a strong, massive and long rostrum, a small, rectangular excisura and a weak, angular antirostrum. the inner face is slightly convex with a broad, long and deepened central sulcus. the cauda is about as long as the ostium with a regularly rounded termination. the ostium is slightly deeper than the cauda and with a distinct anterior opening. the dorsal depression is narrow, small, only above cauda. a ventral line is not present. the outer face is slightly concave and rather smooth except for few marginal furrows at the ventral rim and a longer furrow starting from an incision at the dorsal rim. the central part (nucleus of the otolith) is somewhat depressed and surrounded by indications of growth lines. all rims are sharp. 34 a b c d e f h i j k l m g fig. 11. a–i. genus clupeidarum rectiventralis n. sp. a–c: holotype, selandian, sundkrogen, mguh 26058, × 35. d, e: paratype, selandian, sundkrogen, mguh 26059, × 35. f–i. genus c. aff. rectiventralis, selandian, sundkrogen, mguh 26060–26061, × 35. j–k. arius danicus koken 1891. holotype of arius rotundus roedel 1930, selandian, from erratic boulders of northern germany, pmhub, × 12. l, m. genus salmonidarum acutirostratus n. sp. holotype, selandian, sundkrogen, mguh 26062, × 20. 35 ontogeny and variability. the well-preserved smaller specimen of fig. 11d–e differs in being more compressed. this is likely to represent an ontogenetical effect. two further, smaller and slightly eroded specimens (fig. 11f–i) are even more compressed and show a distinctly convex post-ventral rim. because of this feature, their attribution to the species remains doubtful at present. discussion. this otolith represents the earliest record of the family clupeidae. it shows the typical otolith morphology of this family, i.e. recent representatives of the genera clupea, sardinops and sardinella. nevertheless, genus c. rectiventralis n. sp. cannot with certainty be placed in one of the living genera. the closest fossil resemblance is shared with clupea testis koken 1891 from the oligocene of the north sea basin. this species, however, differs in being more elongated and exhibiting a clear postcaudal furrow on the inner face towards the posterior tip of the otolith. order siluriformes cuvier 1817 family ariidae günther 1864 genus arius cuvier & valenciennes 1840 type species. arius arius cuvier & valenciennes 1840. arius danicus koken 1891 fig. 11j, k 1891 arius danicus koken 1891, p. 81, fig. 1. 1930 arius germanicus koken 1891 – roedel 1930, p. 52. 1930 arius rotundus n. sp. roedel, p. 52, plate 1, fig. 17. material. 67 otoliths; sundkrogen: 12 specimens; vestre gasværk: 50 specimens; gemmas allé: 4 specimens and kongedyb: 1 specimen. in addition, roedel’s holotype of arius rotundus has been inspected and it is redrawn in fig. 11j, k. measurements (roedel’s holotype of arius rotundus) l h t l:h h:t 5.05 3.90 1.55 1.30 2.5 for abbreviations used in the table, see p. 26. description. almost regularly rounded lapilli otoliths except for a mild postdorsal projection. the inner face is convex and smooth; the outer face is flat but with some faint radial furrows. discussion. ariid otoliths are regularly recorded from the early tertiary of europe. so far, arius danicus is the only valid species from the paleocene. order salmoniformes bleeker 1859 suborder salmonoidei bleeker 1859 family salmonidae rafinesque 1815 remarks. recent salmonid fish are mostly freshwater fish that occasionally invade marine environments close to large river mouths. consequently, their occurrence in fossil marine strata is extremely rare and otoliths have so far not been recorded. the species described below thus represents the first and earliest true marine otolith record of this family. genus indet. genus salmonidarum acutirostratus n. sp. fig. 11l, m type locality. sundkrogen. type stratum. selandian, paleocene. derivation of name. acutus and rostratus (latin), refers to the sharp rostrum. holotype. fig. 11l, m, mguh 26062. diagnosis. an otolith, which is very thin, elongated and with a long and sharp rostrum. the ventral rim is shallow; the dorsal rim is with a strong postdorsal angle. the sulcus is wide and long and with only very faint indications of a separation into a shorter ostium and a longer cauda. material. 1 specimen (the holotype). measurements l h t l:h h:t 2.55 1.20 0.20 2.15 6.0 for abbreviations used in the table, see p. 26. description. the otolith is very thin, fragile and elongated. the length:height index is about 2.1. the rostrum is sharp, thin and long. an antirostrum or excisura 36 is not present. the ventral rim is very shallow and gently curved. the dorsal rim is short, highest at the prominent postdorsal angle. the posterior rim is rather regularly curved. all rims are sharp and practically smooth. the inner face is slightly convex, smooth and with a long and wide, slightly supramedian and not much deepened sulcus. differentiation into ostium and cauda is very faint; the colliculi are poorly visible and not separated. the ostium is somewhat shorter than the cauda and faintly widened ventrally. the cauda is straight with a rounded tip close to the posterior rim of the otolith. the dorsal depression is small and faint. a ventral furrow is not present. the outer face is slightly concave and smooth. discussion. the very shallow ventral rim is without any indication of a medioventral angle, which excludes this specimen to represent an argentinid or osmerid genus (see below). this character in combination with the organisation of the sulcus and the strongly developed rostrum is typical for salmonid otoliths. it also distinguishes this otolith from the contemporaneous argentina longirostris, which shares a similarly elongated appearance. a generic identification is not possible at present. genus s. acutirostratus thus represents the first salmonid otolith found in a true marine environment and it also represents the earliest otolith based record of the family. suborder argentinoidei berthelsen 1958 family argentinidae bonaparte 1838 remarks. the family argentinidae is well represented in the paleocene of denmark with four rather common and one rare otolith based species (the latter in open nomenclature) which grow to rather considerable sizes, even when compared to living argentinids. argentinids (and related families) are also rather common in the paleocene and early eocene strata of england, germanyandellesmere island,north-eastcanada (schwarzhans 1986) as well as the maastrichtian and paleocene of bavaria (unpublished data, w. schwarzhans). thus it seems that the cretaceous–paleogene boundary represented an acme zone in the evolution of this family (or suborder). genus protoargentinolithus n. gen. type species. genus berycidarum balticus roedel 1930. derivation of name. combination of proto (greek) and the genus name argentina, referring to the early stratigraphic occurrence and the plesiomorphic features of these otoliths. the ending -lithus is attached to the genus name to indicate that it represents a fossil otolith based genus. diagnosis. a fossil otolith based genus of the family argentinidae with the following characters. the otoliths are elongated, oval in outline and rather thin. the dorsal rim is gently curved, shallow and with or without a postdorsal angle. the ventral rim is more deeply curved, smooth and with a rounded medioventral angle located just slightly in front of the middle. the rostrum is massive but it is not very long. excisura and antirostrum are missing. the sulcus is long, located supramedian, with a long and narrow cauda and a shorter and widened ostium. the dorsal field shows a narrow, but long dorsal depression, whereas the ventral field is completely smooth and without a ventral furrow. discussion. otoliths of the genus protoargentinolithus are very plesiomorphic in appearance, but the prime characters given in the diagnosis above resemble best argentinid otoliths. otoliths can grow to a rather large size (8–9 mm), which is larger than observed in most recent argentinids. in conclusion, protoargentinolithus represents a plesiomorphic extinct member of the argentinidae. distribution. two fossil species are placed in this new genus, both exclusively known from the middle paleocene (selandian) of the north sea basin, i.e. p. balticus from denmark and north-east germany and the new species p. procerus from denmark. protoargentinolithus balticus (roedel 1930) fig. 12a–k 1930 genus berycidarum balticus roedel 1930, p. 62, plate 1, fig. 7. 1930 genus percidarum erraticus roedel 1930 (unfigured paratype only, not holotype – see below). 1930 genus sparidarum gregarius koken 1891 – roedel 1930, p. 66, plate 1, fig. 9. 37 a b c d f g h i j k e fig. 12. protargentinolithus balticus (roedel 1930). a–c, g–k: selandian, sundkrogen, mguh 26063–26069, × 12. d, e: roedel’s holotype, pmhub, × 12. f: paratype of genus percidarum erraticus roedel 1930, pmhub, × 12. 38 material. 624 otoliths from the selandian; sundkrogen: 351 specimens (figured specimens mguh 26063– 26069); vestre gasværk: 252 specimens and kongedyb: 21 specimens. in addition roedel’s holotype is being refigured as fig. 12d, e. a paratype of genus percidarum erraticus is shown in fig. 12f. the specimens identified by roedel as genus sparidarum gregarius koken 1891 have also been inspected, but are not refigured due to the poor preservation. measurements l h t l:h h:t 6.50 3.95 0.95 1.50 4.2 para.* 5.85 3.90 0.85 1.50 4.6 5.35 3.75 n.m. 1.40 ~ 4.40 3.25 0.85 3.8 holo. 2.40 1.74 n.m. 1.35 for abbreviations used in the table, see p. 26. * of g. p. erraticus. description. rather large (up to 6 mm in size) and oval otoliths with gently curved rims without any prominent angles such as a postdorsal angle. the length: height index varies from 1.3 to 1.5.the rostrum is blunt rather short and without antirostrum or excisura. the ventral rim is deeply curved; it is deepest just anterior of themiddle.the inner and outer faces are very smooth; the inner face is slightly convex and the outer face is flat. the sulcus is supramedian with a very long and narrow cauda reaching close to the posterior tip of the otolith and a somewhat widened shorter ostium with a distinct anterior opening. ontogeny and variability. the variations that are observed in the otoliths of this species are moderate. they are mostly concerning details in the development of the dorsal rim and the length:height index. a postdorsal angle is never developed. smaller otoliths below 3.5 mm of length (fig. 11j, k) tend to develop some marginal crenellations along the dorsal and the postventral rims. discussion. protoargentinolithus balticus is easily distinguished from p. procerus n. sp. by its rather small length:height index and the absence of a postdorsal angle. otoliths of argentina erratica are similar in proportions but are always recognised by the prominent and sharp postdorsal angle. also the ostium is shorter and narrower and the inner face is less convex. protargentinolithus balticus is together with coelorhynchus balticus, pteralbula conchaeformis and genus ophidypterus seelandicus the most common otolith based species in the selandian of denmark. protoargentinolithus procerus n. sp. fig. 13a–h type locality. sundkrogen. type stratum. selandian, paleocene. derivation of name. procerus (latin) = elongated, projecting; referring to the elongated shape of these otoliths, which is the main diagnostic feature to distinguish them from the related species p. balticus (see above). holotype. fig. 13d, e, mguh 26070. paratypes. fig. 11a–c, f–h; mguh 26071–26075. diagnosis. elongated, rather fragile and large otoliths (up to 8–9 mm). the length:height index is about 1.7. the ventral rim is rather gently curved, deepest at the middle. the dorsal rim is with a rounded postdorsal angle, which is pointed in large specimens. the cauda is very long and narrow; the ostium is also rather narrow. further material. 146 specimens; sundkrogen: 90 specimens, vestre gasværk: 64 and kongedyb: 6. measurements l h t l:h h:t 8.00 4.90 1.45 1.65 3.4 holo. 4.75 2.90 0.45 1.65 6.5 * ~ 3.80 2.25 n.m. para. ~ 2.50 1.55 n.m. para. for abbreviations used in the table, see p. 26. * not figured, marginally eroded. description.theotolithsareelongatedand fragilemostly of sizes from 4 to 5 mm, but apparently growing up to 8 to 9 mm in length (holotype). the length: height index is ranging from 1.65 to 1.75. the ventral rim is rather regularly and gently curved and deepest at the middle. the dorsal rim is shallow somewhat irregularly ornamented and with very faint and rounded postdorsal angle, which in very large specimens can become pointed. the posterior rim is rounded and 39 a b cd e f g h fig. 13. protargentinolithus procerus n. sp. a–c, f–h: paratypes, selandian, a, b from vestre gasværk and c–h from sundkrogen, mguh 26071–26075, × 12. d, e: holotype, selandian, sundkrogen, mguh 26070, × 12. 40 somewhat shifted dorsally. the rostrum is massive and not very long; in practically all specimens the rostrum is broken. no excisura or antirostrum is present. the dorsal, posterior and postventral rims are finely crenellated in smaller specimens. the inner face is slightly convex and with a slightly supramedian sulcus. it has a marked dorsal depression and a smooth ventral field without ventral furrow. the sulcus is very long and narrow in particular the cauda, which terminates close to the posterior rim of the otolith.the ostium is much shorter and not much widened but somewhat deeper. the outer face is rather smooth and flat. all rims are sharp. ontogeny and variability. p. procerus is the largest argentinid species found in the paleocene of denmark and, judging from the size of its otoliths, one of the largest species of this family at all. like in p. balticus, smaller specimens of p. procerus show finely crenellated dorsal, posterior and postventral rims. very large specimens (fig. 13a–c) exhibit a strong postdorsal angle, which is practically absent in smaller specimens. variability is rather moderate and confined to details in the expression of the dorsal rim. discussion. although in many specimens of p. procerus the rostrum is somewhat damaged, it is obvious in most that the otoliths are considerably more elongated than those of the related p. balticus. small and poorly preserved specimens of p. procerus can possibly be confused with those of the two parallel occurring speciesof thegenusargentina. thespecimensofa. erratica that are slightly less elongated exhibit a deeply curved ventral rim and always a prominent postdorsal angle. those of a. longirostris likewise show the prominent postdorsal angle but have a similar length:height index. however, in this species the tapering caudal tip is connected with the posterior rim by a narrow depression. genus argentina linnaeus 1758 type species. argentina sphyraena linnaeus 1758. argentina erratica (roedel 1930) fig. 14a–l 1930 genus percidarum erraticus roedel 1930, p. 67, plate 1, fig. 11. 1965 primaevomesus tricrenulatus n. sp. stinton, p. 399, plate 30, figs 6, 7, plate 33, fig. 35. ?1966 elops undulatus n. sp. stinton, p. 418, plate 66, fig. 1. 1966 hypomesus pennatus n. sp. stinton, p. 421, plate 66, fig. 6. 1986 argentina pennata stinton 1966 – schwarzhans, p. 788–790, figs 9, 10. material. 207 otoliths from the selandian; sundkrogen: 176 (figured specimens mguh 26080–26083), vestre gasværk: 20, gemmas allé: 1 and kongedyb: 10 (figured specimens mguh 26076–26079). in addition roedel’s holotype has been reviewed and is refigured as fig. 14a, and a valid paratype as fig. 14b, c. measurements l h t l:h h:t 3.65 2.65 0.65 1.40 4.0 3.50 2.40 0.45 1.45 5.3 2.55 1.80 n.m. 1.40 2.15 1.50 n.m. 1.45 1.95 1.30 n.m. 1.50 1.85 1.15 n.m. 1.60 1.50 0.95 n.m. 1.55 for abbreviations used in the table, see p. 26. description. moderately large (up to 4 mm) and thin otoliths with the typical pentagonal outline of argentinid and osmerid otoliths. the five angles are the massive pointed rostrum, the rounded predorsal and the more pronounced postdorsal angles, the rounded angle at the posterior rim and the rounded medioventral angle at the deeply curving ventral rim. excisura and antirostrum are missing. the length:height index ranges from 1.45 to 1.65. the rather flat inner face shows a long and narrow supramedian sulcus. in particular the cauda is long and narrow and about two times as long as the ostium and is reaching far back towards the posterior rim of the otolith. near the tip of the cauda the dorsal crista typically fades away. the ostium is rather short and only slightly widened. the dorsal depression is rather large and marked. the ventral field is smooth sometimes with an indication of a ventral furrow very close to the ventral rim of the otolith. the outer face is nearly flat and smooth. ontogeny and variability. smaller specimens tend to be more irregularly ornamented along the dorsal rim 41 a b c d e f g h i j k l fig. 14. argentina erratica (roedel 1930). a: roedel’s holotype, pmhub, × 20. b, c: roedel’s paratype, pmhub, × 20. d–h: selandian, kongedyb, mguh 26076–26079, × 20. i–l: selandian, sundkrogen, mguh 26080–26083, × 20. 42 than larger ones and in most cases are also slightly more elongated. the variability is mainly confined to the length:height index. discussion. although roedel’s holotype lacks the rostral tip it is still preserved well enough to serve as a holotype. i have no doubt that the specimens, which are from the same formation and the same geographic region and described here, belong to a. erratica. in fact, it is one of the most common species in certain locations of the selandian of denmark (i.e. sundkrogen). the strong postdorsal angle distinguishes these otoliths from the two parallel occurring protoargentinolithus species even in small specimens. its more elongated shape and the postcaudal connection to the posterior rim characterise argentina longirostris, described below. argentina erratica also seems to be a common species in the late paleocene (thanetian) and early eocene (london clay) of the north sea basin from where it has been described by stinton (1965, 1966) under several names (see synonymy list). argentina pennata, which is recorded from theearlyeoceneof theellesmere island, arctic ocean (schwarzhans 1986) very likely also represents the same species. in fact, characters of all the above mentioned records are so similar that they likely represent a single species. argentina longirostris n. sp. fig. 15a–g type locality. sundkrogen. type stratum. selandian, paleocene. derivation of name. longirostris (latin), referring to the long rostrum of the species. holotype. fig. 15c, d, mguh 26084. paratypes. fig. 15a, b, e–g, mguh 26085–26089. diagnosis. thin, elongated and fragile otoliths with a length:height index of 1.7 to 1.9. the rostrum is sharp and long; no antirostrum or excisura are present. the postdorsal angle is prominent, the predorsal angle is almost absent. the sulcus is supramedian, long and narrow; the caudal tip is reaching very close to the posterior tip of the otolith and is connected with it by a small depression. further material. 60 specimens; sundkrogen: 56; vestre gasværk: 4. measurements l h t l:h h:t 5.45 3.05 n.m. 1.80 para. 4.00 2.05 0.4 1.95 5.1 holo. 3.25 1.85 n.m. 1.75 para. 2.25 1.25 n.m. 1.80 para. 1.35 0.80 n.m. 1.70 para. for abbreviations used in the table, see p. 26. description. the otoliths are very elongated, thin and fragile reaching to about 6 mm in length. the ventral rim is smooth, rather shallow and regularly curved, deepest near the middle. the dorsal rim is somewhat undulated with a prominent postdorsal angle and a very faint predorsal angle. the posterior tip is blunt, markedly shifted towards the dorsal. the anterior tip has a long, fragile and sharp rostrum. the antirostrum and excisura are missing. the inner face is slightly convex, smooth and with a supramedian sulcus. the sulcus is very long and narrow, deep and indistinctly divided into a shorter ostium and a much longer cauda. the caudal tip reaches very close to the posterior rim of the otolith and is connected with it via a small depression. the ostium is very slightly widened and opened anteriorly. the dorsal depression is narrow and not very distinct. the ventral field is smooth, sometimes with a very faint indication of a ventral furrow close to the ventral rim of the otolith. the outer face is nearly flat and smooth. all rims are sharp. ontogeny. smaller specimens, i.e. less than 3 mm in length are slightly more compressed than the larger ones. otherwise the variability is mainly confined to minor differences in the expression of the dorsal rim and the length:height index. discussion. this is a typical representative of the genus argentina, which like all recent argentina otoliths exhibits the connection of the caudal tip with the posterior rim via a narrow depression. small specimens can be confused with either a. erratica or protoargentinolithus procerus, particularly when the specimens are fragmented. however, complete specimens differ from those two species in the more elongated shape and the postcaudal connection to the posterior rim. 43 b c d e f g h i j a fig. 15. a–g. argentina longirostris n. sp. a, b, e–g: paratypes, selandian, sundkrogen, mguh 26085–26089, × 20. c, d: holotype, selandian, sundkrogen, mguh 26084, × 20. h–j. genus argentinidarum sp. selandian, sundkrogen, mguh 26090–26092, × 20. 44 genus indet. genus argentinidarum sp. fig. 15h–j material. sundkrogen: 6 specimens (mguh 26090– 26092) and kongedyb: 1 specimen. measurements l h t l:h h:t 2.35 1.15 n.m. 2.05 for abbreviations used in the table, see p. 26. description. otoliths are small, 2.5 to 3 mm, very elongated and rather thin. the length:height index is about 2.0. the dorsal and ventral rims are shallow with blunt median angles. the rostrum and posterior tip is pointed. the cauda is very long, more than two times of the ostium and with a broadly rounded tip. the inner face is almost flat with some marginal furrows to the posterior on the ventral field and occasionally an indication of a ventral furrow is present and not too close to the ventral rim of the otolith. discussion. these otoliths clearly represent an undescribed species probably of some kind of argentinid or related family. however, none of the specimens is preserved well enough to serve as holotype. order aulopiformes rosen 1973 family aulopidae cope 1872 genus aulopus cloquet 1816 type species. salmo filamentosus bloch 1792. aulopus tortus n. sp. fig. 16a–f type locality. fakse quarry, se sjælland. type stratum. fakse coral limestone, danian, paleocene (leg. a. rasmussen). derivation of name. from tortus (latin) = twisted; referring to the torsion of the otolith along the horizontal axis. holotype. fig. 16b–e, mguh 26093. paratypes. fig. 16a, f, mguh 26094–26095. diagnosis. massive and elongated otoliths that are somewhat twisted along the horizontal axis and with a length:height index between 2.2 and 2.4. the rostrum is massive and long; the posterior tip is projecting dorsally. the sulcus isnarrowanddeep; thecauda is straight and longer than the ostium and somewhat inclined downwards. the ostium is dorsally widened. further material. 1 specimen. measurements l h t l:h h:t 2.85 1.25 n.m. 2.30 para. 2.65 1.20 0.60 2.20 2.0 holo. for abbreviations used in the table, see p. 26. description. otoliths are small, probably not exceeding 3.5 mm, very elongated and rather massive. the ventral rim is shallow, gently curved and smooth. the dorsal rim is short and nearly flat. the rostrum is rather long and massive. an antirostrum is not present and the excisura is only incipient. the posterior tip is strongly projecting and pointed dorsally. the inner face is distinctly twisted along the horizontal axis and has a long, narrow and deep sulcus. the cauda is straight and much longer than the ostium. it is somewhat inclined downwards and terminates with a rounded tip close to the postventral rim. the ostium is considerably deepened somewhat widened to the dorsal side and with a distinct anterior opening. the dorsal depression is very shallow. the ventral furrow is feeble and runs very close to the ventral rim of the otolith. the outer face is convex and smooth, also depicting the twist along the horizontal axis. the rims are rather thick. discussion. the characters given in the diagnosis easily recognise a. tortus n. sp. the torsion of the otolith along its horizontal axis and the downward inclination of the cauda are typical characters for aulopid otoliths and distinguish these otoliths from chlorophthalmus postangulatus, which occurs simultaneously. 45 a bc d e f g h i j k l m n o p fig. 16. a–f. aulopus tortus n. sp. a, f: paratypes, danian, fakse quarry, mguh 26094–26095, × 20. b–e: holotype, danian, fakse quarry, mguh 26093, × 20. g–p. chlorophthalmus postangulatus nolf & dockery 1993. g–i, k–p: selandian, sundkrogen, mguh 26097–26103, × 20. j: danian, fakse quarry, mguh 26096, × 20. 46 family chlorophthalmidae jordan 1923 genus chlorophthalmus bonaparte 1840 type species. chlorophthalmus agassizi bonaparte 1840. chlorophthalmus postangulatus nolf & dockery 1993 fig. 16g–p 1993 genus chlorophthalmidarum postangulatus nolf & dockery, pp. 28–30, plate 2, figs 1, 2. ?1993 genus argentinoideorum sculptissimus n. sp. nolf & dockery, p. 28, plate 2, fig. 7. material. 169 otoliths: fakse quarry: 3 (danian; figured specimen mguh 26097) and 166 (selandian); sundkrogen: 46 (figured specimens mguh 26097– 26103); vestre gasværk: 113; gemmas allé: 4 and kongedyb: 3. measurements l h t l:h h:t 3.85 1.85 0.85 2.10 2.2 3.05 1.60 n.m. 1.90 2.60 1.45 0.55 1.80 2.6 2.05 1.05 n.m. 1.95 for abbreviations used in the table, see p. 26. description. moderately elongated to elongated, small and massive otoliths. the ventral rim is gently and shallowly curved and smooth. the dorsal rim is with midand postdorsal angles of variable intensity often irregularly undulated. the rostrum is short, massive and pointed. an antirostrum or excisura are not present. the posterior tip is rounded or somewhat projecting dorsally. the inner face is slightly convex with a supramedian, shallow, narrow and long sulcus. the cauda is very narrow and bends slightly downwards to the posterior terminating at a moderate distance from the posterior rim. the ostium is short and somewhat widened, especially to the dorsal side. the dorsal depression is rather wide and distinct. the ventral field is smooth and without a ventral furrow. the outer face is slightly convex and smooth. the rims are moderately sharp. ontogeny and variability. the variability in this species is quite large and eye-catching. it mostly concerns differences in thedevelopmentof thedorsal rim.whereas in some species there is a strong postdorsal angle and virtually no mediodorsal one others show (in addition) a rather prominent mediodorsal angle. these differences also find their expression in the variability of the length:height index. it seems that in smaller specimens of about 3 mm and less the mediodorsal angle is more pronounced whereas in large specimens it is mostly reduced. also small specimens show a stronger marginal crenellation. the large variability found in the chlorophthalmid species is also highlighted by the fact that nolf & dockery (1993) have described a separate species under the name of genus argentinoideorum sculptissimus based on a unique specimen, which likely represents only an abnormal specimen of c. postangulatus. discussion. chlorophthalmid otoliths form a common faunal element in late cretaceous and early paleogene deposits. their otoliths depict a rather generalised morphology and together with the considerable variability individual species are often difficult to be distinguished, particularly so when dealing with smaller specimens of 3 mm length and less. the large specimens described here very closely resemble those described from the paleocene of alabama (usa) by nolf & dockery (1993) and are interpreted to represent the same species. it is also known from the paleocene of bavaria (unpublished data, w. schwarzhans). another species, commonly described from the early eocene of the london basin as synodus davisi (frost 1925), also represents a chlorophthalmus species. differences to c. postangulatus seem very small if at all valid at species level. i have not inspected the type specimen of c. davisi so the possibility of placing c. postangulatus into synonymy is omitted until a more widely review of the various fossil species of this genus has been performed. order myctophiformes regan 1911 family myctophidae gill 1893 genus indet. genus myctophidarum schnetleri n. sp. fig. 17a–f, j type locality. sundkrogen, excavation 1920. type stratum. selandian, paleocene. derivation of name. after k. ingemann schnetler. 47 a b c d e f gh i j holotype. fig. 17a, b, mguh 26104. paratypes. fig. 17c–f, j, mguh 26105–26109. diagnosis. elongated rather massive otoliths with a smooth ventral rim, a pronounced postdorsal angle and a blunt and massive rostrum but without antirostrum or excisura. the sulcus is wide and long; the ostium is somewhat deepened; the cauda is longer than the ostium and nearly equally wide. the caudal colliculum has a sharp ventral margin resembling an incipient caudal pseudocolliculum as it is characteristic for otoliths of all living myctophids. a broad, distinct ventral furrow is present near the ventral rim of the otolith. material. sundkrogen: 9 specimens; vestre gasværk: 2 specimens. measurements l h t l:h h:t 2.25 1.20 n.m. 1.85 para. 2.10 1.15 0.4 1.85 2.9 holo. 2.10 1.25 n.m. 1.70 para. 1.75 1.00 n.m. 1.75 para. for abbreviations used in the table, see p. 26. description. otoliths are rather small, probably not exceeding 2.5 mm, elongated, robust and with a length: height index of 1.7 to 1.9. the ventral rim is shallow, gently curved and smooth. the dorsal rim is almost flat, sometimes undulating and with a distinct postdorsal angle at its end. the anterior rim is with a broad, massive and blunt rostrum, but no antirostrum or excisura is present. the posterior rim is bluntly rounded or cut vertically. the inner face is almost flat with a broad, shallow fig. 17. a–f, j. genus myctophidarum schnetleri n. sp. a, b: holotype, selandian, sundkrogen, mguh 26104, × 20. c–f, j: paratypes, selandian, sundkrogen, mguh 26105–26109, × 20. g–i. genus myctophidarum sp. selandian, vestre gasværk, mguh 26110, × 20. 48 and long central sulcus. the ostium is anteriorly opened and somewhat deepened; it is shorter than the cauda and not widened. the cauda is longer and shallower than the ostium but of similar width and with somewhat upward bend termination not far from the posterior tip of the otolith. the caudal colliculum has a sharp ventral margin resembling the caudal pseudocolliculum observed in all recent representatives of the family. the dorsal depression is rather large and distinct. the ventral furrow is broad, distinct and situated close to the ventral rim of the otolith. the outer face is flat to slightly convex and almost smooth. the rims are moderately sharp. variability. all otoliths known at present are of about the same size. variations are restricted to the ornamentation of the rims and the expression of the posterior rim. discussion. genus m. schnetleri probably represents a new fossil genus, but more material should be inspected before a formal decision is made. the main difference from the otoliths of the living myctophid genera is the lack of a separated caudal pseudocolliculum. however, from the character status in genus m. schnetleri it can be interpreted how that peculiar feature, which is so characteristic for myctophid otoliths, has developed. eokrefftia schwarzhans 1984 is a ‘modern’ myctophid with a separated pseudocolliculum and is already known from the paleocene of south australia. genus indet. genus myctophidarum sp. fig. 17g–i material. 3 otoliths from the selandian; vestre gasværk: 1 well preserved specimen (mguh 26110) and sundkrogen: 2 eroded specimens. measurements l h t l:h h:t 2.90 2.10 0.75 1.40 2.8 for abbreviations used in the table, see p. 26. description. small and massive otoliths with a nearly flat inner face and a distinctly convex outer face. all rims are gently curved and the ventral rim is rather shallow. the rostrum is massive and very pronounced. the ostium and cauda are about equal in length, the ostium somewhat widened; the caudal colliculum has a distinct ventral crest (incipient caudal pseudocolliculum). the ventral furrow is distinct, narrow and close to the ventral rim; the dorsal depression is wide and shallow. discussion. these otoliths closely resemble the maastrichtian and paleocene undescribed specimens from bavaria, germany, where they represent a common faunal element. allocation with the myctophidae is tentative. order percopsiformes berg 1940 family indet. remarks. the order percopsiformes is a small group of fish. three living families are known and all are restricted to the freshwater environment in north america. the order is regarded among the most primitive living representatives of the paracanthopterygii. this view is also supported by otolith investigations. otolithmorphologyshowsbasic resemblancewith those of the orders batrachoidiformes and ophidiiformes, but differs from gadiformes in the lack of a homosulcoid sulcus with a pseudobiostial sulcus opening. the small otoliths described here in many ways resemble certain living percopsiform otoliths like those of the monogeneric north american freshwater families percopsidae and aphredoderidae. the sulcus organisation and the general appearance of those otoliths could well be interpreted as that of a primitive paracanthopterygian close to the percopsiformes. however, the few data available at present do not allow for a more precise identification. in view of the different living habitus of the recent percopsiformes the allocation of the otoliths remains tentative. genus ?percopsiformorum enigmaticus n. sp. fig. 18a–g type locality. sundkrogen. type stratum. lellinge grønsand, selandian. derivation of name. from enigma (greek); referring to the enigmatic allocation of this species. holotype. fig. 18b, c, mguh 26111. paratypes. fig. 18a, d–g, mguh 26112–26115. 49 a b c d e f g fig. 18. genus ?percopsiformorum enigmaticus n. sp. a–d, g: paratypes, selandian, sundkrogen, mguh 26112–26115, × 35. e, f: holotype, selandian, sundkrogen, mguh 26111; e: × 35; f: × 20. diagnosis. small elongated otoliths that rarely exceed 3mmin length.the rostrumand the excisura are marked. the sulcus has an ostial opening; the ostium is slightly longer but not wider than the cauda. the colliculi are well marked; the caudal colliculum is antero-ventrally somewhat reduced giving way to an incipient, indistinct crest (?pseudocolliculum). the inner face is rather flat. further material. vestre gasværk: 1 specimen. measurements l h t l:h h:t 3.20 1.85 n.m. 1.75 para. 2.85 1.60 0.50 1.80 3.2 holo. 2.10 1.25 n.m. 1.70 para. 1.80 1.15 n.m. 1.55 para. 1.15 0.75 0.25 1.55 3.0 para. for abbreviations used in the table, see p. 26. description. otoliths are small (size up to 3.5–4 mm) 50 and elongated. the ventral rim is shallow, smooth and gently curved; the dorsal rim is likewise shallow and with broadly rounded medioand postdorsal angles. the anterior tip has a sharp, pointed rostrum, a sharp excisura and a moderate to faint antirostrum. the posterior tip is pointed, but it is more rounded than the anterior tip and is located slightly inframedian. all rims are smooth or slightly undulating. the inner face is nearly flat with a large central sulcus. the sulcus is shallow with a clear ostial opening; the sulcus is divided into a slightly longer ostium and a shorter cauda. both the cauda and the ostium bend slightly upwards and are of similar width. the colliculi are well marked, the caudal colliculum is somewhat reduced towards the antero-ventral giving way to an incipient and indistinct crest (?pseudocolliculum). the dorsal depression is small and indistinct and without a ventral furrow. the outer face is slightly convex and rather smooth. the rims are thin. ontogeny. small specimens are more compressed than the large ones. the rims are more irregularly ornamented in the small ones. in general otoliths of less than 2 mm of length apparently are juveniles and do not show all the valid diagnostic features. specimens between 2.5 to 3 mm in length are morphologically well defined. discussion. as stated above in the introduction to percopsiformes these otoliths exhibit a number of unusual characters that on the one hand make them easy to recognise and on the other hand make the systematic-phylogenetic interpretation very difficult. the systematic position for genus ?percopsiformorum enigmaticus therefore remains tentative. order gadiformes goodrich 1909 remarks. the systematic of gadiforms has been a field of extensive study and phylogenetic (re-) evaluations, and thus in recent years has resulted in several new andalternativecladistic concepts.althoughmostof these concepts do not integrate morphological analyses of otoliths it is apparent that many of the new proposals for gadiform classification reflect otolith findings much better than the previous ones. i have in the following used the classification of markle (1989) which mostly fits well with otolith findings. the fishes of the gadiform families ranicipitidae, merlucciidae, lotidae and gadidae form a common faunal element in the temperate and cool seas of the northern hemisphere, both recent and in the fossil record. their otoliths are well known from the tertiary of the north sea basin, particularly since oligocene times. they are also common in the paleocene and the early eocene (london basin), but are entirely missing from the middle and late eocene of the north sea basin. this phenomenon coincides with a warm temperature maximum during a short lived connection of the southern north sea basin with the atlantic ocean that brought in the warm water adapted fauna that is so well known from the eocene of the belgium basin. cretaceous gadiforms are not known from otoliths or skeletons. the paleocene fish fauna of the north sea basin, however, is already quite rich in gadiform species as based on otoliths from various gadiform families. most of these forms represent rather ‘primitive’ genera or groups and the most relevant groups of the recent fauna of the area are already present. the otoliths from the selandian of denmark have contributed particularly to the record (see following list). ranicipitidae: raniceps hermani nolf 1978 from the selandian of denmark and the thanetian of belgium. merlucciidae: palaeogadus sinangulatus n. sp. from the selandian of denmark. euclichthyidae: archaemacruroides ornatus stinton 1965 (for synonymy see nolf 1978), a plesiomorphic form (and fossil otolith based genus) possibly related to the recent euclichthys, described from the thanetian of england and belgium and a yet undescribed second species of archaemacruroides from the paleocene of bavaria, germany. lotidae:gadomorpholithusponderosus (koken1885), which is a very plesiomorphic lotid otolith (and fossil otolith based genus) and molva palaeomorpha n. sp.; both are from the selandian of denmark. gadidae: protocolliolus amorphus n. sp. (fossil otolith based genus) near trisopterus from the selandian of denmark. macrouridae: coelorhynchus balticus (koken 1885) from the selandian of denmark and north-east germany and two more species from the selandian of denmark, i.e. hymenocephalus rosenkrantzi n. sp. and coryphaenoides amager n. sp. in fact, very small larval and indeterminable gadiform otoliths represent the most common element in the fish fauna of the selandian of denmark. 51 a b c d e family ranicipitidae markle 1989 genus raniceps oken 1817 type species. blennius raninus linnaeus 1758. raniceps hermani nolf 1978 fig. 19a–e 1978 raniceps hermani nolf 1978, p. 225, plate 1, fig. 5. material. vestre gasværk: 3 specimens (mguh 26116– 26118), selandian. measurements l h t l:h h:t 6.35 2.65 n.m. 2.40 4.85 2.10 1.00 2.30 2.1 2.95 1.45 0.75 2.05 1.9 for abbreviations used in the table, see p. 26. description. the otoliths are very elongated and with a length:height index well above 2.0. the dorsal and ventral rims are shallow, gently curved and without prominent angles. the anterior tip is broadly rounded and the posterior tip is pointed. the postdorsal rim is slightly crenellated; the other rims are smooth in adults. the inner face is flat to slightly convex and smooth; it has a large, wide and shallow sulcus. the ostium is slightly shorter than the cauda and is also slightly narrower. the colliculi are large, oval, flat and in level with the surface of the inner face. the caudal colliculum is reduced towards its posterior tip and the ostial colliculum is reduced towards its anterior tip. this in combination with the very narrow collum results in the sulcusmorphology typical of thegenus raniceps. the ventral furrow is indistinct and close to the ventral rim. theouter facehas a feeble precentral umbo and some radial furrows are crossing the surface. ontogeny. the three specimens available represent a fig. 19. raniceps hermani nolf 1978. selandian, vestre gasværk, mguh 26116–26118; a, c–e: × 20; b: × 12. 52 rather complete ontogenetic succession. the smallest juvenile specimen is remarkable for its more compressed appearance and the more intense ornamentation of the otolith rims. discussion. raniceps hermani nolf 1978 was first described from the thanetian of belgium (sands of orp, now interpreted as selandian, np4–np5; nolf, personal communication 1999). family merlucciidae gill 1884 genus palaeogadus rath 1859 type species. palaeogadus troscheli rath 1859 (= nemopteryx crassus agassiz 1843). palaeogadus sinangulatus n. sp. fig. 20a–i type locality. sundkrogen. type stratum. lellinge greensand, selandian, paleocene. derivation of name. sine (latin) = without and angulus (latin) = angle; referring to the rounded predorsal rim in adults. holotype. fig. 20d, e, mguh 26119. paratypes. vestre gasværk: fig. 20a–c, mguh 26121; fig. 20f, mguh 26120; gemmas allé: fig. 20g, h, mguh 21622; sundkrogen: fig. 20i, mguh 26123. diagnosis. elongated large otoliths with the subtriangular outline typical for the otoliths of genus palaeogadus. the predorsal angle is broad and not prominent and is getting completely rounded in adult specimens. the anterior and posterior tips are rounded. the inner face is convex. thecollum is ratherwide andwithout a pseudocolliculum. the ventral furrow is distinct and very close to the ventral rim. further material. 173 otoliths; sundkrogen: 44 specimens, vestre gasværk: 113 specimens, gemmas allé: 5 specimens and kongedyb: 11 specimens. measurements l h t l:h h:t 6.45 2.70 1.25 2.40 2.2 para. 5.45 2.45 1.30 2.25 1.9 holo. 5.10 2.30 n.m. 2.20 para. 4.65 2.05 1.05 2.25 1.9 para. 3.85 1.85 0.95 2.10 1.9 para. for abbreviations used in the table, see p. 26. description. elongated otoliths having a roughly triangular outline and mostly smooth rims. the ventral rim is shallow and gently curved towards the anterior and almost straight to the posterior. the dorsal rim has a rather feeble predorsal angle near to the anterior tip of the otolith; in large specimens it is getting rounded and reduced. the posterior tip is somewhat pointed. the anterior tip is bluntly pointed below the ostium. the size of the otoliths reaches about 7 mm. the inner face is slightly convex and with a large median homosulcoid sulcus. the ostium is shorter than cauda. both are filled with distinct colliculi that in small specimens are somewhat reduced in size outwards to the otolith margins. the collum is rather wide and without central pseudocolliculum. the ventral furrow is distinct and very close to the ventral rim. the dorsal depression is very narrow and indistinct. the outer face is flat to slightly convex. it is rather smooth with an indistinct pre-central umbo and occasional radial furrows. the rims are smooth to moderately sharp and sometimes slightly undulated, which is particularly true for the dorsal rim. ontogeny and variability. the ontogenetic allometric growth in otoliths of this species is remarkable. the main changes occur late in the ontogeny and in otoliths of the size of about 5 mm. at about this size the predorsal angle becomes reduced, the anterior tip of the otolith is more gently rounded and the ostial and caudal colliculi now completely fill the ostium and cauda respectively (except for the rather wide collum). also the large otoliths show a tendency to develop some indistinct marginal undulation or even ornamentation, which in a way is a reversal of the usually observed ontogenetic trend. fortunately, this species is common, particularly at the vestre gasværk location, so a complete ontogenetic succession is well represented. variability on the other hand is rather restricted. it is mainly the length:height index and the expression of the predorsal angle that vary to a certain degree. 53 a b c d e f g h i fig. 20. palaeogadus sinangulatus n. sp. a–c: paratype, selandian, vestre gasværk, mguh 26121; a: × 20; b, c: × 12. d, e: holotype, selandian, sundkrogen, mguh 26119; d: × 12; e: × 20. f, i: paratypes, selandian, sundkrogen, mguh 26120–26123, × 20. g, h: paratype, selandian, gemmas allé, mguh 26122, × 20. 54 discussion. p. sinangulatus closely resembles p. trigonus, which has been described from the thanetian (late paleocene) of the london basin (stinton 1965). the single known specimen of p. trigonus is poorly preserved and due its small size seemingly did not develop all the pertinent diagnostically valid characters. i therefore follow nolf (1985) in rejecting p. trigonus as a valid species. p. sinangulatus is based on a good sequence of well-preserved and diagnostically mature specimens and thus represents a well-defined species. the situation should be further evaluated when more palaeogadus specimens of thanetian age become available. as noted above, smaller specimens of p. sinangulatus resemble other typical species of the genus palaeogadus, an extinct genus well known both based on otoliths and skeletons from the early tertiary of northern europe. the large specimens have a lot in common withcertainplesiomorphicgadidgenerasuchascolliolus or protocolliolus, due to the ontogenetic alteration of details of the outline (see above).morphologically, these otoliths are somewhat intermediate between the merlucciid genus palaeogadus and early gadids. at the same time, however, true gadid otoliths occur as well. family lotidae bonaparte 1832 genus gadomorpholithus n. gen. type species. genus gadidarum ponderosus koken 1885. derivation of name. combination of the genus name gadus and morpho (greek) = form, referring to the gadid ‘look-alike’ plesiomorphic features of these otoliths. the ending -lithus is attached to the genus name to indicate it representing a fossil otolith based genus. diagnosis. a fossil otolith based genus of the family lotidae with the following characters: the otoliths are elongated and moderately thin and with pointed anterior and posterior tips. the dorsal rim shows a prominent, broadly rounded predorsal angle and a very feeble postdorsal angle. the ventral rim is gently curved and finely crenellated. the sulcus is long, typical homosulcoid and pseudobiostial in organisation and located slightly supramedian. the cauda is somewhat longer than the ostium; both are nearly completely filled with well-defined colliculi. a central pseudocolliculum is missing. the dorsal field shows a narrow but long dorsal depression. the ventral field exhibits a clear cut and long ventral furrow not very close to the ventral rim of the otolith. above the ventral furrow the ventral field is smooth, whereas some marginal furrows occur near to the rim. the outer face is flat to concave and without an umbo, but with intense ornamentation along the dorsal and ventral rims. otoliths of this genus can grow to a rather large size (about 10 mm). discussion. otoliths of genus gadomorpholithus combine plesiomorphic characters of the merlucciidae with apomorphic characters of the lotidae and gadidae. plesiomorphic characters are the lack of a central pseudocolliculum and the broad predorsal angle. apomorphic lotid / gadid characters are the form of the anterior tip of the otolith and the large colliculi. the strongly convex ventral field, the development of the ventral line, the postdorsal angle and the concave outer face are more typical for lotids. all in all, i assume that gadomorpholithus represents a very primitive and basal phylogenetic member of the family lotidae. distribution. a single fossil species (g. ponderosus) is from the middle paleocene (selandian) of denmark. gadomorpholithus ponderosus (koken 1885) fig. 21a–m 1885 gadidarum ponderosus koken 1885, p. 113, plate 5, fig. 24. material. 87 otoliths from the selandian of denmark; sundkrogen: 19 specimens (figured specimens mguh 26124–26127); vestre gasværk: 63 specimens (figured specimens mguh 26128–26130), gemmas allé: 4 specimens and kongedyb: 1 specimen. measurements l h t l:h h:t 8.00 3.45 1.55 2.30 2.2 6.55 3.05 1.30 2.15 2.3 4.30 1.90 0.80 2.25 2.4 for abbreviations used in the table, see p. 26. description. see diagnosis for the genus (monospecific genus). ontogeny. specimens of about 6 to 7 mm and larger (including koken’s holotype) represent truly adults (fig. 21a–h). smaller ones in the order of 4 to 6 mm (fig. 55 a b c d e f g h i j k l m fig. 21. gadomorpholithus ponderosus (koken 1885). a–c, h–j, l, m: selandian, sundkrogen, mguh 26124–26127; a, h–j, l, m: × 20; b, c: × 12. d–g, k: selandian, vestre gasværk, mguh 26128–26130; d, g, k: × 12; e, f: × 8. 56 a b c d f g h e 21i–m) are more generalised in several characters of the outline, the sulcus and the curvature of the inner face. also the ornamentation of the otolith rims and the outer face is more extensive. remarks. among the many larval gadiform otoliths of less than 3 mm of size there may be several specimens that may belong to this species. the small sized specimens however have not developed valid diagnostic characters that allow distinction from other parallel occurring gadiforms. genus molva lesueur 1819 type species. gadus molva linnaeus 1758. molva palaeomorpha n. sp. fig. 22a–h type locality. vestre gasværk. type stratum. lellinge greensand, selandian, paleocene. derivation of name. from palaeo (greek) = old and morpho (greek) = form; referring to the early occurrence of this typical lotid otoliths. fig. 22. molva palaeomorpha n. sp. a, b: holotype, selandian, vestre gasværk, mguh 26131; a: × 20; b: × 12. c, d: paratypes, vestre gasværk, mguh 26132–26133, × 20. e–h: paratypes, sundkrogen, mguh 26134–26135, × 20. 57 holotype. fig. 22a, b, mguh 26131. paratypes. fig. 22c, f, topoand stratotypes, mguh 26132–26133; fig. 22d, e, g, h sundkrogen, mguh 26134–26135. diagnosis. thin and elongated otoliths with a slightly convex inner and a concave outer face. the dorsal rim is straight and the ventral rim is very shallow. the anterior tip is inframedian and the posterior tip is supramedian. the colliculi are about equal in size, tapering and pointed towards the narrow collum and widening to the outer margins. further material. sundkrogen: 2 specimens; vestre gasværk: 1 specimen. measurements l h t l:h h:t 6.05 2.45 0.90 2.45 2.7 holo. 2.95 1.45 0.45 2.05 3.2 para. 1.95 1.00 0.40 1.95 2.5 para. for abbreviations used in the table, see p. 26. description. otoliths are rather thin and elongated with the typical parallelogram-like outline. the dorsal rim is straight with rounded predorsal angle; the ventral rim is gently curved and shallow. the anterior is tip blunt and inframedian. the posterior tip is blunt and supramedian. all rims are smooth to slightly undulate. otolith size is 6 mm and more. the inner face is slightly convex with a supramedian, homosulcoid and rather shallow and narrow sulcus. the colliculi are well defined, about equal in size and tapering and pointed towards the narrow collum but widening to the outer margins. the ventral furrow is feeble and rather close to the ventral rim. the dorsal depression is rather small and moderately deepened. the outer face is slightly concave and with little ornamentation. the rims are sharp. ontogeny. the holotype is the only surely adult specimen available. the otoliths of this species can be reasonably well recognised down to a size of about 2 to 2.5 mm due to the very specific outline and sulcus morphology. there is, however, a gap in the ontogenetic sequence. next to the holotype of about 6 mm of length the next smaller specimen is only about half the size (3.2 mm). this and other specimens of the size mainly differ in the more rounded posterior and anterior tips and the much smaller length:height index. the smallest specimen of about 2 mm is also the most compressed. also it is thicker with a shallow precentral umbo on the outer face. discussion. m. palaeomorpha is a typical representative of the family lotidae and can be placed in the genus molva with a reasonable certainty. gadomorpholithus ponderosus, which occurs simultaneously differs in being more robust (thicker) and in the proportions of the anterior-ventral and in the posteriordorsal rims. family gadidae rafinesque 1810 genus protocolliolus gaemers 1976 type species. gadus eocenicus frost 1931. protocolliolus amorphus n. sp. fig. 23a–j type locality. sundkrogen, excavation 1920. type stratum. lellinge grønsand, selandian, paleocene. derivation of name. amorphus (latin) = amorphous, referring to the generalised appearance of the otolith. holotype. fig. 23a–c, mguh 26136. paratypes. sundkrogen, fig. 23d, e, mguh 26137; gemmas allé, fig. 23f–j, mguh 26138–26140. diagnosis. robust, thick and elongated otoliths with a regularly rounded or bluntly pointed anterior rim and a pointed posterior tip. the dorsal rim is regularly curved with an indistinct predorsal angle; the ventral rim is smooth and deepest to the anterior of the middle. the inner face is markedly convex, rather smooth and with a ventral furrow very close to the ventral rim. the sulcus is moderately shallow, long, wide, homosulcoid and pseudobiostial. the cauda is longer than the ostium and both are completely filled with the colliculi. a central pseudocolliculum is not present. the outer face is convex with very little ornamentation. further material. 103 specimens from the selandian; sundkrogen: 1 specimen, vestre gasværk: 98 specimens; gemmas allé: 4 specimens. 58 a b c d f g h ij e fig. 23. protocolliolus amorphus n. sp. a–c: holotype, selandian, sundkrogen, mguh 26136; a, c: × 12; b: × 20. d, e: paratype, vestre gasværk, mguh 26137; d: × 12; e: × 20. f–j: paratypes, gemmas allé, mguh 26138–26140; f, h–j: × 20; g: × 12. 59 measurements l h t l:h h:t 6.00 3.50 1.65 1.70 2.1 holo. 4.70 2.60 1.30 1.80 2.0 para. 4.25 2.40 1.15 1.75 2.1 para. 3.50 1.95 n.m. 1.80 para. 1.65 0.95 0.45 1.75 2.1 para. for abbreviations used in the table, see p. 26. description. otoliths are medium in size, growing up to 6 mm, massive and robust and elongated in shape. the outline is typically ‘seed-like’ with a high, blunt or regularly rounded anterior rim and a moderately pointed posterior tip. the dorsal rim is gently curved and somewhat undulating with an indistinct and broadly rounded predorsal angle. the ventral rim is likewise gently curved, smooth and deepest to the anterior of the middle. the inner face is markedly convex in horizontal and vertical directions, smooth with a moderately shallow, wide, long and slightly supramedian sulcus, which is typically homosulcoid and pseudobiostial in organisation. the cauda is longer but not wider than the ostium. both are completely filled with the colliculi. the ostial colliculum is somewhat deepened at its centre and so is the caudal colliculum in its anterior portion. the collum is moderately wide without central pseudocolliculum. the dorsal field has a long and shallow dorsal depression. the ventral field is smooth except for a distinct ventral furrow very close to the ventral rim of the otolith. the outer face is convex and with a faint precentral umbo and little ornamentation. all rims are rather thick. ontogeny and variability. even relatively small specimens of p. amorphus, i.e. at a size of less than 2 mm may be recognised by their compressed and thick appearance (fig. 23i, j). however, these otoliths have a very generalised morphology. also the degree of ornamentation decreases with growth, whereas the thickness and length:height ratio remain rather stable. variability seems to be rather restricted to details of the dorsal rim and the degree of ornamentation. discussion. p. amorphus is only the second species described in the genus protocolliolus, which differs from the younger fossil otolith based genus colliolus mainly in the absence of a central pseudocolliculum. in this character it resembles the living trisopterus (fossil evidence since middle oligocene). p. amorphus is rare in most localities of the selandian of denmark compared to other gadiforms, but it is quite common at vestre gasværk. this species is always easily recognised by its massive appearance and the other characters given in the diagnosis. p. eocenicus from the early eocene of the london basin is quite similar, but it is more elongated, thinner and with a more pronounced predorsal angle. family macrouridae jordan & evermann 1898 the family macrouridae is typical for deeper marine shelf environments and continental slopes. in the fossil record their otoliths occur mostly in pelagic and hemipelagic environments. so far, nezumia lindsayi schwarzhans 1984 from the paleocene of south australia is the first paleocene record. now, the paleocene of denmark has yielded three different species: coelorhynchus balticus (koken 1885), which previously was regarded as a merlucciidae, hymenocephalus rosenkrantzi and coryphaenoides amager. amazingly, all three species are already typical macrourids that can be assigned to living macrourid genera with good confidence. this is not an expected finding, because so far there is no support by the skeleton record. furthermore, this indicates that the origin of the family should reach further back in geological time. genus coelorhynchus giorna 1809 type species. lepidoleprus coelorhynchus risso 1810. coelorhynchus balticus (koken 1885) fig. 24a–l 1885 merluccius balticus koken, p. 113, plate 5, fig. 22. 1930 merluccius schmitti n. sp. roedel, p. 54, plate 1, fig. 1. ?1930 merluccius globulosus n. sp. roedel, pp. 54–55, plate 1, fig. 2 (juvenile specimen). material. 474 otoliths from the selandian of denmark; sundkrogen: 87 specimens (figured specimens mguh 26142–26145); vestre gasværk: 371 specimens (figured specimen mguh 26141); gemmas allé: 7 specimens, and kongedyb: 9 specimens. in addition the type specimens of merluccius balticus from koken (1885) and merluccius schmitti and m. 60 fig. 24. coelorhynchus balticus (koken 1885). a–c: selandian, vestre gasværk, mguh 26141; a: × 12; b, c: × 8. e: holotype of merluccius schmitti roedel 1930, selandian, from erratic boulders of northern germany, pmhub, × 12. d, f–l: selandian, sundkrogen, mguh 26142–26145, × 12. a b c d e f g h i j k l 61 globulosus from roedel 1930 have been inspected. the type specimen of m. schmitti is illustrated in fig. 24e. measurements l h t l:h h:t 9.80 5.10 2.05 1.90 2.50 6.75 3.55 1.50 1.90 2.35 holo.* 6.00 3.10 1.10 1.95 2.80 2.80 1.55 0.60 1.80 2.60 for abbreviations used in the table, see p. 26. * of merluccius schmitti. description. elongated and thin otoliths with a triangular outline. the otolith size reaches up to 10 mm. the anterior tip is rounded somewhat pronounced below the ostium and the posterior tip is pointed. the ventral rim is gently curved almost flat in the central portion and deepest anteriorly. the dorsal rim is with a prominent predorsal angle; preand postdorsal parts of the dorsal rim are almost straight and downwards inclined. all rims are intensely ornamented. the inner face is convex and smooth with deeply invading furrows from the ornamentation of the rims with a rather narrow and somewhat deepened sulcus situated distinctly supramedian. the cauda is almost two times as long as the ostium; both are filled with oval and anteriorly and posteriorly rounded and somewhat deepened colliculi. the collum is narrow and without central pseudocolliculum. the dorsal depression is small and well marked ventrally. the ventral furrow is feeble or absent and close to the ventral rim. the outer face is flat, intensely and deeply ornamented. ontogeny and variability. the largest otoliths, which are about 6.5 mm and more, are less ornamented than the smaller ones. otoliths of less than 4–5 mm tend to become more compressed and generalised in character and specimens of 3–2.5 mm and less cannot be identified to a species particularly when the predorsal angle becomes rounded. i assume that many if not most of the larval unidentifiable gadiform otoliths described below probably represent this species, which is the most common gadiform found in the selandian of denmark. the variability is moderate in otoliths of similar sizes. it is confined to the intensity of the ornamentation and details of the outline. discussion. roedel’s holotype of merluccius schmitti is perfectly preserved and would serve well as a holotype for this species. koken’s holotype of merluccius balticus is also a fairly large specimen, but it lacks the posterior third of the otolith and is also slightly eroded. nevertheless, the typical form of the sulcus and the anterior part of the outline are well enough preserved to allow identification, particularly since now sufficient newly collected material is available from the same locality for the redefinition of the species. c. balticus is a typical representative of the genus coelorhynchus and does not differ much from otoliths of recent species of this genus. typical characters are the outline with the pronounced predorsal angle and the large sulcus with its large and long colliculi. genus coryphaenoides gunnerus 1765 type species. coryphaenoides rupestris gunnerus 1765. coryphaenoides amager n. sp. fig. 25a–i type locality. vestre gasværk. type stratum. lellinge grønsand, selandian, paleocene. derivation of name. after amager, the island south of copenhagen, where the gemmas allé locality was exposed. holotype. fig. 25e–h, mguh 26146. paratypes. fig. 25a–d, topoand stratotype, mguh 26147–26148; fig. 25i, sundkrogen, mguh 26149. diagnosis. oval to elongated and rather thin otoliths. the anterior tip is rounded, the posterior tip is somewhat pointed; the ventral rim is distinctly pronounced anteriorly; the dorsal rim is gently curved and without a predorsal angle. the sulcus is narrow. the colliculi are small and the caudal colliculum is reduced to the posterior. the collum is narrow and without pseudocolliculum. further material. vestre gasværk: 5 specimens; sundkrogen: 1 specimen. 62 a b c d e g h f i fig. 25. coryphaenoides amager n. sp. a–d: paratypes, selandian, vestre gasværk, mguh 26147–26148, × 12. e–h: holotype, selandian, vestre gasværk, mguh 26146, × 12. i: paratype, selandian, sundkrogen, mguh 26149, × 12. measurements l h t l:h h:t 3.50 1.10 3.2 para. 6.90 3.35 1.50 2.05 2.2 para. 4.40 2.35 0.85 1.90 2.8 holo. 3.95 2.05 n.m. 1.95 para. for abbreviations used in the table, see p. 26. description. otoliths are rather thin, elongated and oval in outline. the size is up to 7–8 mm. the ventral rim is gently and regularly curved and distinctly pronounced anteriorly; the dorsal rim is also gently curved and without predorsal angle. the anterior rim is broadly rounded with an inframedian tip; the posterior tip is rounded or somewhat pointed. all rims are finely crenellated becoming smooth in adults. the inner face is slightly convex with a narrow, slightly supramedian and shallow sulcus. the cauda is not much longer than the ostium. the caudal colliculum is posteriorly reduced and therefore it is not much larger than the ostial colliculum. the collum is narrow and without pseudocolliculum. the ventral furrow is feeble and close to the ventral rim. the dorsal depression is indistinct and small. marginal zones near the otolith rims are with numerous radial furrows originating from the marginal crenellations. the outer face is flat to slightly convex and intensely ornamented. the rims are sharp. ontogeny. most specimens available of this species are rather small subadults (including the holotype). only two of the paratypes (fig. 25a–d) are from truly adult specimens. of these two one is complete but marginally somewhat eroded whereas the other is well pre63 a b c e f g d fig. 26. hymenocephalus rosenkrantzi n. sp. a, b, f, g: paratypes, selandian, kongedyb, mguh 26151–26154, × 20. c–e: holotype, selandian, kongedyb, mguh 26150; c: × 20; d, e: × 12. served but lacking the posterior tip. anyhow, they both show that the main ontogenetic trend to be expected and concerning the drastic reduction of the marginal ornamentation and also the ornamentation on the outer face. discussion. this seems to be a typical representative of the genus coryphaenoides. the regular outline without a predorsal angle is diagnostic for species of this genus in combination with the posteriorly reduced cauda and caudal colliculum and the narrow collum without a pseudocolliculum, which also distinguishes it from parallel occurring gadiform species. genus hymenocephalus giglioli 1884 type species. hymenocephalus italicus giglioli 1884. hymenocephalus rosenkrantzi n. sp. fig. 26a–g 64 type locality. kongedyb i. type stratum. lellinge grønsand, selandian, paleocene. derivation of name. the species is named after the late professor in palaeontology a. rosenkrantz. holotype. fig. 26c–e, mguh 26150. paratypes. fig. 26a, b, f, g, topoand stratotype, mguh 26151–26154. diagnosis. moderately small and compressed otoliths with a distinct predorsal angle. the inner face is rather flat. the cauda is not much longer than the ostium; the caudal colliculum is posteriorly reduced and thus about equally long as the ostial colliculum. the collum is moderately wide with a small pseudocolliculum. the ventral furrow is distinct and close to the ventral rim. further material. 10 otoliths; vestre gasværk: 8 specimens; kongedyb: 2 specimens. measurements l h t l:h h:t 4.95 2.95 n.m. 1.65 para. 4.25 2.80 n.m. 1.50 para. 3.90 2.55 1.00 1.55 2.5 holo. 3.30 2.20 n.m. 1.50 para. 1.70 1.10 n.m. 1.55 para. for abbreviations used in the table, see p. 26. description. the otoliths are compressed, massive and rather small reaching to about 5–6 mm in size. the ventral rim is smooth, deeply and gently curved and distinctly pronounced to the anterior. the dorsal rim is shallower except for the prominent predorsal lobe. the anterior tip is bluntly rounded and ventrally pronounced; the posterior tip is pointed. all rims are rather smooth, but the dorsal rim has some irregular undulations. the inner face is rather flat, smooth and with a slightly supramedian, wide and slightly deepened sulcus. the cauda is not much longer than the ostium; the caudal colliculum is somewhat reduced to the posa b c d e f g h fig. 27. gadiformes sp. juveniles. a–c: danian, fakse quarry, mguh 26155, × 20. d: holotype of merluccius nanus roedel 1930, erratic boulders of northern germany, pmhub, × 20. e–h: selandian, sundkrogen, mguh 26156, × 35. 65 terior and thus is not longer than the ostial colliculum. the collum is moderately wide with a small pseudocolliculum. the ventral furrow is close to the ventral rim; the dorsal depression is large, but is not distinct. the outer face is more convex than the inner face, without a distinct umbo and rather smooth with few feeble radial furrows. the rims are moderately thick. ontogeny and variability. the largest otolith at hand is just slightly more elongated than the other otoliths but otherwise it is very similar. otoliths of 3 mm and more in size seem to be diagnostically mature. smaller ones can not always be identified because they have a verygeneralisedmorphology. sometimes thesmall specimens can be identified by their compressed appearance. the overall small size of these otoliths is well in line with recent species of this genus. discussion. these otoliths resemble those of certain recent species of the genus hymenocephalus in the overall characters. recent otoliths of this genus are known for their very large pseudocolliculum and sometimes the fusion of the colliculi. there are also species with a more ‘normal’-sized pseudocolliculum and wellseparated colliculi (see schwarzhans 1981a). h. rosenkrantzi resembles the latter. i tentatively regard this fossil species as a very primitive member of the genus probably close to its phylogenetic origin. gadiformes spp. fig. 27a–h material. 734 otoliths; 4 otoliths from the danian of fakse (fig. 27a–c, mguh 26155); 730 otoliths from the selandian, 447 from sundkrogen (fig. 27e, f, mguh 26156; fig. 27g, h); 246 from vestre gasværk; 7 from gemmas allé; the holotype of merluccius nanus roedel 1930 (fig. 27d). remarks. gadiformes spp. comprises juvenile and not identifiable specimens. the most common otoliths found in the selandian of denmark and especially at sundkrogen are small gadiform otoliths of sizes between 1.5 to 2.5 mm. the danian location at fakse bears similar otoliths although they are less common. morphologically these small otoliths have avery generalised appearance so they cannot be identified at the specific level. they probably represent juvenile or larval fishes. it is even impossible to judge whether they represent merlucciid, gadid, macrourid or some other gadiform species of which large and diagnostically valuable otoliths have been described above. it is possible that many of them represent coelorhynchus balticus, which is the most common gadiform species (based on large otoliths), from these locations. more likely though is that they simply represent juveniles and larval specimens of several of the recorded gadiform species. the specimen figured in fig. 27e, f may, however, represent a juvenile of a yet undescribed gadiform species. this tiny otolith is remarkable for its symmetrical appearance (outline and sulcus) and the combination of a strongly convex inner face and a flat to concave outer face. because of the taxonomical uncertainties i propose to leave these otoliths in open nomenclature. likewise, roedel’s holotype of merluccius nanus (fig. 27d) represents such a small gadiform specimen and i cannot recommend using this species name. order ophidiiformes berg 1937 family ophidiidae rafinesque 1810 genus indet. genus ophidiidarum seelandicus (koken 1885) fig. 28a–k 1885 trachinus seelandicus koken, p. 115, plate 5, fig. 25. 1930 merluccius latisculptatus n. sp. roedel, p. 56, plate 1, fig. 3. ?1965 ophidypterus retusus n. sp. stinton, p. 408, plate 31, fig. 20, plate 33, fig. 44. material. 696 otoliths from the selandian of denmark; sundkrogen: 130 specimens (figured specimens mguh 26157–26162); vestre gasværk: 538 specimens (figured specimen mguh 26163); gemmas allé: 18 specimens and from kongedybet: 10 specimens. roedel’s fragmented holotype of merluccius latisculptatus is refigured in fig. 28c. measurements l h t l:h h:t 4.25 2.35 0.95 1.80 2.5 3.35 1.80 0.75 1.85 2.4 2.45 1.35 0.60 1.80 2.3 1.85 1.05 n.m. 1.75 for abbreviations used in the table, see p. 26. 66 a b c d f g h i j k e fig. 28. genus ophidiidarum seelandicus (koken 1885). a, b, f–k: selandian, sundkrogen, mguh 26157–26162, × 20. c: holotype of merluccius latisculptatus roedel 1930, erratic boulders of northern germany, pmhub, × 35. d, e: selandian, vestre gasværk, mguh 26163, × 20. 67 short description. the otoliths are oval to moderately elongated and thin. the otolith is reaching up to 5 mm in size. the anterior tip is blunt and the posterior tip is pointed at the middle. the ventral and dorsal rims are gently curved and somewhat undulating. the dorsal rim has weak preand postdorsal angles; the highest point of the otolith is to the anterior of the mid-part. the inner face is slightly convex rather smooth and with a long, median to slightly supramedian sulcus. the sulcus is closed anteriorly not far from the anterior tip of the otolith and it is clearly subdivided into a somewhat longer, shallow and ventrally widened ostium and a shorter, deeper and ventrally narrowed cauda; the cauda turns slightly downwards towards its tip. ostial and caudal colliculi are well marked. the dorsal depression is rather small; it is shallow and indistinct. the ventral furrow is distinct and runs close to the ventral rim. it starts from the ostial tip and terminates close to the caudal tip of the sulcus. the outer face is rather flat and smooth or slightly ornamented. the rims are moderately sharp. ontogeny and variability. smaller otoliths, i.e. below 3 mm in size, differ from the adults in four characters. these are (1) the less pointed posterior tip of the otolith, (2) the more pronounced postdorsal angle, (3) the more flat ventral rim and (4) the more thick appearance of the rims. these characters are typical of juvenile ophidiid otoliths. in principle such juvenile ophidiid otoliths are not diagnostically mature. they can only be specifically identified in the presence of a good ontogenetic sequence. variability is less prominent, particularly among the adults and is restricted to details of the outline of the otoliths. discussion. stinton (1965) established the fossil otolith based genus ophidypterus with ophidypterus retusus as the holotype. the single otolith is a juvenile and less than 2.5 mm in length and was obtained from the late paleocene (thanetian) strata in england. the holotype must be regarded as a not diagnostically mature specimen and until further material has been found the validity of this species (and the genus) must be in doubt. it does, however, resemble small otoliths of o. seelandicus and may well fall within the variability of that species. nolf (1978) described similar but poorly preserved otoliths as ampheristus sp. from the thanetian strata in belgium. these specimens may represent another species, possibly of the genus hoplobrotula that is characterised by having a short cauda. also, stinton (1977) from the thanetian in england has described a true hoplobrotula species, i.e. h. protensa. this species is characterised by having a short cauda and an elongated otolith shape. family bythitidae gill 1861 genus bidenichthys barnard 1934 type species. bidenichthys capensis barnard 1934. bidenichthys lapierrei (nolf 1978) fig. 29a–j 1978 ogilbia lapierrei nolf, p. 226, plate 2, figs 2, 3. material. 71 otoliths from the danian of fakse (figured specimens mguh 26164–26171); 6 otoliths from the selandian of denmark; 2 from sundkrogen, 2 from vestre gasværk and 2 from kongedyb. measurements l h t l:h h:t 5.05 2.45 1.20 2.05 2.0 * 3.75 1.95 0.95 1.90 2.0 * 3.60 1.65 0.65 2.15 2.5 † 3.15 1.65 n.m. 1.90 * 2.35 1.25 n.m. 1.90 * 2.05 1.10 n.m. 1.85 † 1.95 0.95 0.55 2.05 1.7 † for abbreviations used in the table, see p. 26. * danian specimen. † selandian specimen. description. small, compact and elongated otoliths that reach a size to about 5 mm. the anterior tip is bluntly pointed at about the middle. the posterior tip has a massive blunt projection, which dorsally is marked by a distinct indentation. the ventral rim is flat and smooth, deepest to the anterior of the mid-part of the specimen. the dorsal rim is likewise flat and smooth and nearly straight betweenthe roundedpreand postdorsal angles. the inner face is slightly convex with a moderately long central sulcus that terminates at some distance from the anterior and posterior tips of the otolith. the ostium is about twice as wide and more than twice as long as the cauda. the colliculi are well marked and distinctly separated and deepened. the dorsal depression is large but with indistinct margins. the ventral 68 a b c d e f g h i j k l m n o fig. 29. a–j. bidenichthys lapierrei (nolf 1978). danian, fakse quarry, mguh 26164–26171, × 12. k–o. genus bythitidarum rasmussenae n. sp. k, l, o: paratypes, danien, fakse quarry, mguh 26173–26175, × 20. m, n: holotype, danian, fakse quarry, mguh 26172, × 20. 69 furrow is feeble and long; it is running very close to the ventral rim of the otolith. the outer face is slightly convex and smooth. all rims are moderately thick. ontogeny and variability. ontogenetic changes and the variability are limited. they are restricted to details of the outline and minor variations in the proportions. small specimens below 2.5 to 3.0 mm are somewhat generalised in outline and less diagnostically valid. discussion. the genus bidenichthys shows one of the most primitive otolith patterns within the family bythitidae. primitive characters are the clear separation into the ostium and cauda and the deepening of the sulcus in particular of the cauda. most other recent bythitid otoliths show more or less fused colliculi and a flat sulcus. three endemic species of the genus bidenichthys are known to day from near shore environments and tidal pools in south africa and new zealand. this distribution pattern is interpreted as a secondary endemism. in the fossil record the otoliths of the genus bidenichthys are quite common in some late cretaceous and early tertiary sediments. distinction of the various species as discussed in the following is mainly based on differences in the proportion of the otolith and the sulcus and certain characters of the outline. the earliest fossil record is b. crepidatus (voigt 1926) from the campanian of northern germany and the maastrichtian of bavaria, germany (unpublished data, w. schwarzhans). these otoliths lack the posterior projection and postdorsal indentation and dorsal and ventral rims are more regularly curved. b. midwayensis (nolf & dockery 1993) is very similar and was originally described as belonging to another bythitid genus (i.e. dinematichthys) from the paleocene in alabama, usa. it mainly differs in being more elongated. another very similar yet undescribed species was found in the paleocene of bavaria, germany (unpublished data, w. schwarzhans). this species differs in having a more pronounced ventral rim to the posterior and a more strongly reduced anterior portion of the ostium resulting in a relatively large distance between the ostial tip and the anterior tip of the otolith. in the eocene of the north sea basin b. sagittalis (frost 1934) has mainly been recorded from england. these otoliths are distinctly more elongated and with a length:height index above 2.20. the youngest fossil record from europe is b. boscheineni schwarzhans 1994 from the late oligocene of northern germany, where it occurs in a facies similar to the rock pool environment of the recent species. the otoliths of this species are easily recognised by their pronounced ventral rim to the posterior and by the sulcus proportions, which are much reduced in length. b. lapierrei (nolf 1978) was originally described as ogilbia lapierrei from the thanetian of belgium. all characters are so similar to the specimens from the danian of fakse that i have no doubt placing them in the same species. b. lapierrei has rarely been found in the selandian, which may be due to environmental reasons. genus indet. genus bythitidarum rasmussenae n. sp. fig. 29k–o derivation of name. after mrs. a. rasmussen (fakse), who collected most of the otoliths from the fakse quarry. holotype. fig. 29m, n, mguh 26172. type locality. fakse quarry. type stratum. soft coral limestone, early paleocene, danian. paratypes. fig. 29k, l, o, topoand stratotypes, mguh 26173–26175. diagnosis. small compact and rather compressed otoliths. the anterior tip is bluntly pointed and the posterior tip has a massive projection and a distinct postdorsal indentation. the ventral rim is rather deeply curved. the sulcus is short, wide, deep and with completely fused colliculi. a feeble indentation at the ventral rim of the sulcus marks a previous separation into a large ostium and a small cauda. further material. 1 specimen from fakse quarry. measurements l h t l:h h:t 3.05 1.80 n.m. 1.70 para. 2.70 1.55 n.m. 1.75 para. 2.65 1.55 0.80 1.70 1.9 holo. 2.25 1.30 n.m. 1.75 para. for abbreviations used in the table, see p. 26. 70 description. the otoliths are small, not exceeding 3 mm, thick, compact and rather compressed. the anterior tip is bluntly pointed. the posterior tip is a broad, massive projection and dorsally marked by a distinct postdorsal indentation. the ventral and dorsal rims are smooth and gently, but rather deeply curved without prominent angles. the inner face is moderately convex with a wide not very long but considerably deepened central sulcus. the sulcus terminates at some distance from the anterior and the posterior tips of the otolith. a single and completely fused colliculum is well marked. a feeble incurvation at the ventral rim of the sulcus marks a ‘former’ separation of the sulcus into a large and wide ostium and a small and narrow cauda. the dorsal depression is rather large and deep. the ventral furrow is feeble or absent and mostly visible only near the anterior tip of the otolith. the outer face is more strongly convex and smooth, which results in a very massive appearance of the otoliths. the rims are rather thick. ontogeny and variability. this species is smaller than the simultaneously occurring bidenichthys lapierrei (see above). specimens from 2.5 to 3.0 mm are well developed and diagnostically mature. only the smallest specimen with a size of about 2.3 mm is somewhat generalised in outline. variability seems to be restricted to details of the outline. discussion. otoliths of genus b. rasmussenae are well distinguished from bidenichthys lapierrei by their more compressed and compact appearance and the completely fused single colliculum. the presence of fused colliculi precludes allocation of genus b. rasmussenae as a species of the genus bidenichthys, although in many other characters it resembles bidenichthys otoliths quite well. the tendency of fusing the colliculi is well known from several lineages within the bythitidae including the subfamily dinematichthyinae to which this species most likely belongs. it can not directly be related to any of the recent dinematichthyin genera with fused colliculi and probably it represents an extinct lineage. order lampridiformes regan 1909 family veliferidae bleeker 1859 genus indet. genus veliferidarum harderi n. sp. fig. 30a–g, fig. 31a, c, g type locality. sundkrogen, excavation 1920. type stratum. middle paleocene, selandian. derivation of name. after p. harder, who collected most of the otoliths from the sundkrogen locality. holotype. fig. 30e–g, mguh 26176 paratypes. fig. 30a–d, fig. 31a, c, vestre gasværk, selandian, mguh 26177–26180; fig. 31d, sundkrogen, selandian, mguh 26181; fig. 31b (aff. harderi), fakse, danian, mguh 26182. diagnosis. delicate, compressed and rounded otoliths with a massive but not very long rostrum. the dorsal rim is crenellated or undulated. the preventral rim is undulated and the postventral rim is smooth and deeply curved. the sulcus is typically heterosulcoid. the ostium is open, short and ventrally widened to the anterior; the cauda is long, narrow and almost straight. further material. 74 specimens from the selandian; 35 from sundkrogen, 38 from vestre gasværk and 1 from kongedyb. measurements l h t l:h h:t 5.50 5.15 n.m. 1.05 para.* 4.65 4.40 1.15 1.05 3.8 para.* 3.65 3.50 0.65 1.05 5.5 holo. 2.55 2.35 n.m. 1.10 para.† 2.95 2.45 n.m. 1.20 para.‡ for abbreviations used in the table, see p. 26. * from vestre gasværk. † from sundkrogen. ‡ from fakse. description. otoliths are thin, fragile, getting more robust with growth, compressed and with a rounded outline. the otolith size is up to 6 mm. otolith rims are rounded with feeble midand postdorsal angles, sometimes with a pre-ventral angle and a short pointed posterior tip (in adult specimens only). the rostrum is 71 a b c d e f g fig. 30. genus veliferidarum harderi n. sp. a–d: paratypes, selandian, vestre gasværk, mguh 26177–26178; a, c: × 20; b, d: × 12. e–g: holotype, selandian, sundkrogen, mguh 26176, × 20. 72 massive, short and blunt. the excisura is not very deep; the antirostrum is weak. the entire dorsal rim is intensely crenellated or undulated and the anterior part of the ventral rim is sometimes finely crenellated and posteriorly roughly undulated. the inner face is markedly convex, particularly in the horizontal direction. the sulcus is heterosulcoid in organisation, situated slightly supramedian and somewhat deepened.the ostium is short and anteriorly open; ventrally it is considerably widened. the cauda is about 1.5 times as long as the ostium; it is narrow, almost straight and with a ventrally pointed tip close to the posterior rim of the otolith. the colliculi have rather indistinct margins. the dorsal field has a small, narrow depression and with some radial furrows originating from the crenellations of the dorsal rim. the ventral field is rather smooth and without a ventral furrow. the outer face is concave with some radial furrows near the rims. all rims are sharp. ontogeny and variability. at a size between 4 and 5 mm the otoliths of genus v. harderi change morphology and one could be tempted to regard them as different species in cases where the ontogenetic succession is not well represented. the large otoliths (more than 4.5 mm; fig. 30a–d) are slightly more compressed because of a deeper ventral rim, which is distinctly more median. the dorsal rim becomes less strongly ornamented and with a little pointed posterior tip. the ventral rim is also smoother to the posterior and develops a serrated ornamentation to the anterior. in general the larger otoliths are more robust than the smaller ones. the single specimen from the danian of fakse (fig. 31b) is just slightly more elongated than the specimens from the selandian. it is possible that this specimen represents yet another species, but more material has to beevaluatedbefore adecisioncanbe made.thus this specimen is referred to as genus v. aff. harderi. discussion. an undescribed species from the maastrichtian of bavaria, germany (unpublished data, w. schwarzhans) is less compressed than genus v. harderi but otherwise it is very similar. the correlation with otoliths of the two recent species of the genus velifer is not entirely certain and thus the placement of the fossil species in the family remains tentative. the large specimens of genus v. harderi also show resemblance to otoliths of the zeiform family antigoniidae. order zeiformes regan 1909 family indet. genus indet. genus zeiformorum janni n. sp. fig. 31d–f type locality. gemmas allé. type stratum. middle paleocene, selandian. derivation of name. after my son jan-philipp. holotype. fig. 31d–f, mguh 26183. diagnosis. a small compressed and rounded otolith with a blunt massive rostrum and a broadly rounded posterior rim. the ventral rim is deepest at its middle. the sulcus is long, deepened and anteriorly open; posteriorly it is almost reaching the posterior tip of the otolith. the colliculi are small, widely separated and somewhat deepened. the ventral furrow is short and indistinct only below and close to the collum. measurements l h t l:h h:t 1.90 1.85 0.45 1.05 4.0 for abbreviations used in the table, see p. 26. description. the unique holotype is a very small otolith, i.e. just slightly less than 2 mm, but it is well preserved. the otolith is compressed, very high and with a rounded outline. the dorsal, ventral and posterior rims are deep, regularly curved, almost smooth and without any prominent angles. the anterior rim shows a massive short rostrum, a broad, not very deep excisura and a rather distinct antirostrum. the inner face is slightly convex with a somewhat deepened central sulcus, which is anteriorly opened to the excisura and posteriorly reaches very close to the posterior rim of the otolith. the cauda is just slightly shorter than the ostium; both are separated by a long collum.ostial and caudal colliculi are widely separated and considerably deepened.the ostial colliculum opens towards the excisura. the dorsal field is smooth and without a marked depression; the ventral field is smooth except for an indistinct and short ventral furrow that is close to and only below the collum. the outer face is almost flat and smooth. the rims are sharp. 73 a b c d e f g fig. 31. a, c, g. genus veliferidarum harderi n. sp. a, c: paratypes, selandian, vestre gasværk, mguh 26179–26180, × 20. g: paratype, selandian, sundkrogen, mguh 26181, × 20. b. genus veliferidarum aff. harderi danian, fakse quarry, mguh 26182, × 20. d–f. genus zeiformorum janni n. sp. holotype, selandian, gemmas allé, mguh 26183, × 35. 74 a b c d e f g h fig. 32. hoplostethus lacinatus (koken 1885). a, b: selandian, vestre gasværk, mguh 26184; a: × 20; b: × 12. c–f, h: selandian, sundkrogen, mguh 26185–26188, × 20. g: selandian, gemmas allé, mguh 26189, × 20. 75 discussion. the morphology of this otolith is intermediate between the two main types of morphologies that are found in otoliths of the zeiformes. one morphology type comprises zenionidae and parazenidae in which the colliculi are flat, the ventral furrow is still at some distance from the sulcus, the posterior tip of the otolith is rounded and the caudal tip is closed. the second morphology type comprises the families grammicolepididae, oreosomatidae, cyttidae and zeidae. it is characterised by having a ridge like elevation of the very narrow portion located between the ventral furrow and the sulcus and the biostial opening of the sulcus, i.e. the caudal tip opens into an excisuralike incision at the posterior rim. zeiform otoliths are mostly small. therefore, it is concluded that the unique holotype of genus z. janni is diagnostically mature despite of its small size. other undescribed and similar species have been found in the maastrichtian and paleocene of bavaria. amanses sulcifer described by stinton (1966) from the early eocene of the london basin may also be a related species. these otoliths are even more compressed and with a length:height index that is less than 1; the posterior rim is almost straight and vertical and with the cauda almost opening to it. order beryciformes regan 1909 suborder berycoidei regan 1909 family berycidae lowe 1843 genus hoplostethus cuvier 1829 type species. hoplostethus mediterraneus cuvier 1829. hoplostethus lacinatus (koken 1885) fig. 32a–h 1885 genus apogonidarum lacinatus koken, p. 117, plate 5, fig. 26. material. 103 otoliths from the selandian; 35 from sundkrogen (figured specimens mguh 26185–26188), 64 from vestre gasværk (figured specimen mguh 26184), 1 from gemmas allé (mguh 26189) and 3 from kongedyb. measurements l h t l:h h:t 8.00 7.00 2.35 1.15 3.0 * 6.65 5.65 1.70 1.20 3.3 3.65 3.35 1.00 1.10 3.3 2.65 2.50 n.m. 1.05 1.20 1.35 n.m. 0.90 for abbreviations used in the table, see p. 26. * koken’s holotype. description. these otoliths are moderately large and massive and grow up to 7–8 mm. they are compressed, high with rounded midand postdorsal angles, sharp sometimes pointed preand postventral angles and an angularposterior tip that is located high above the cauda. the rostrum is short, massive and blunt; the excisura is sharp, narrow and not very deep; the antirostrum is feeble. the dorsal and sometimes also the posterior rims are crenellated and the ventral rim is smooth. the inner face is rather flat, particularly in the vertical direction and with a large, somewhat deepened, slightly supramedian sulcus. the ostium is anteriorly opened, ventrally it is considerably widened and shorter than cauda. the ostial colliculum is distinct and somewhat deepened. the cauda is narrower and longer than the ostium and turns upwards towards its tip, where it reaches rather close to the posterior rim of the otolith. the caudal colliculum has a characteristic ridge-like ventral margin. the dorsal depression is large, wide and deep; some radial furrows that originate from the ornamentation of the dorsal rim cross the dorsal depression. the ventral field is smooth and has a very indistinct ventral furrowsituated close to theventral rim. the outer face is convex; it has a broad, smooth subcentral umbo and some radial furrows dorsally. ontogeny and variability. ontogenetic changes in berycid otoliths are quite remarkable (see schwarzhans 1981a), and this is also the case in h. lacinatus. koken’s holotype is a large truly adult specimen, larger than most of the specimens available to me (about 8 mm). the largest figured specimen (fig. 32a, b) has a dorsal rim, which is reduced in the height and ornamentation. otherwise, it resembles the specimens of intermediate sizes from 2.5 to 4 mm (fig. 32c–f). specimens of less than 2.5 to 3 mm in size (fig. 32g, h) must be regarded as juveniles without all of the diagnostically valid characters. the smaller the specimens the stronger the ornamentation and the higher the dorsal field. the smallest specimens are below 1.5 mm and are remarkable for their length:height index of 76 a b c d e f g h i j fig. 33. centroberyx integer (koken 1885). a, e–h: selandian, vestre gasværk, mguh 26190–26192; a, e, g, h: × 12; f: × 20. b–d, j: selandian, sundkrogen, mguh 26193– 26195, × 12. i: danian, fakse quarry (coll. rasmussen), × 12. 77 about 1 or less. because of this it must once again be stressed to only attempt to specifically identify berycid otoliths in the presence of a suitable ontogenetic sequence including truly adult forms. variability on the other hand is less prominent, restricted to the proportions of the otolith and the sulcus and details of the outline. here again, variations are strongest in small specimens. discussion. h. lacinatus is a well-known and easily recognised otolith in the paleocene of denmark. from the simultaneously occurring centroberyx integer and centroberyx fragilis it is distinguished by its more compressed and massive appearance and the distinctive hexagonal outline with a rather strong and massive rostrum. genus centroberyx gill 1862 type species. beryx lineatus cuvier & valenciennes 1829. centroberyx integer (koken 1885) fig. 33a–j 1885 genus apogonidarum integer koken 1885, p. 114, plate 5, fig. 27. 1978 trachichthodes integer koken 1885 – nolf 1978, p. 228. material. 55 otoliths; 2 from the danian of fakse; 53 from the selandian, 27 from sundkrogen (figured specimens mguh 26193–26195) and 26 from vestre gasværk (figured specimens mguh 26190–26192). measurements l h t l:h h:t 8.15 6.15 n.m. 1.30 5.35 4.30 1.30 1.25 3.3 4.60 4.00 n.m. 1.15 3.80 3.00 1.15 1.25 2.6 2.55 2.30 0.85 1.10 2.7 1.40 1.20 n.m. 1.15 for abbreviations used in the table, see p. 26. description. massive and rather thick otoliths with rounded outline growing in size to about 9 mm. the dorsal rim is rather regularly curved, but always with a pronounced, rounded postdorsal angle. the ventral rim is deep with broadly rounded preand postdorsal angles located at large distance from each other; the medioventral rim is in between straight and horizontal. the rostrum is very short and blunt; the excisura and the antirostrum are very feeble or missing. the posterior rim is blunt and dorsally pronounced. the dorsal and the postventral rims may be somewhat crenellated, particularly in juveniles. the inner face is moderately convex to almost flat in vertical direction. the sulcus is supramedian, very long, wide, rather shallow and anteriorly open; posteriorly it terminates close to the posterior tip of the otolith. the ostium is ventrally considerably widened and long, almost as long as the cauda. the ostial colliculum is distinct. the cauda is narrower and longer than the ostium, turning upward towards its tip and reaching rather close to the posterior rim of the otolith. the caudal colliculum has a typical ridge like ventral margin. the dorsal depression is wide and large. the ventral field is smooth sometimes with an indistinct ventral furrow very close to the ventral rim. the outer face is slightly convex, smooth in adults, intensely ornamented with radial furrows in juveniles. the rims are moderately sharp. ontogeny and variability. ontogenetic changes in this species are moderate, but very much in line with those described by schwarzhans (1981a) for centroberyx and also the ones described above for hoplostethus lacinatus. specimens of about 4 to 5 mm can be regarded as diagnostically mature. smaller specimens (fig. 33i, j) tend to show a somewhat higher dorsal rim, which results in a lesser length:height index. also, the marginal ornamentation is stronger. variability is moderate and confined to details of the outline and the ornamentation. discussion. koken’s holotype is a large, well-preserved specimen. it is very similar to the one of fig. 33c, d. small otoliths of c. integer could be confused with small specimens of hoplostethus lacinatus however they never become as compressed. specimens of both species of more than 4 mm are always easily distinguished. with larger specimens centroberyx fragilis (see below) bears more resemblance, but is always distinguished by the more delicate, thin appearance, the more narrow, tapering caudaand the shapeof the ventral rim. centroberyx otoliths have commonly been reported from the late cretaceous and the early tertiary of europe. together with species of related genera, they form a common faunal element in many teleost faunas in the world of that time. in northern europe, c. inte78 ger and c. fragilis are two common species in the paleocene and one common species in the eocene and oligocene is c. subrotundus (koken 1884) (described under many synonyms; see schwarzhans 1981a). c. subrotundus rather closely resembles c. fragilis , but it is more compressed. there is at least one further undescribed species from thepaleocene ofbavaria (unpublisheddata,w.schwarzhans).c. teumeri (voigt 1926) (unpublished data, w. schwarzhans) is a widespread species in the maastrichtian of northern germany and bavaria. it resembles c. integer, but it is more rectangular in outline caused by the flat dorsal rim. other similar species for instance have been reported from the paleocene of the us gulf coast, i.e. genus berycidarum stringeri nolf & dockery 1993, from the eocene of new zealand, i.e. c. pulcher (schwarzhans 1981a) and egregioberyx erectus schwarzhans 1981a and egregioberyx sphaeroides (stinton 1958) from the eocene of australia. centroberyx fragilis n. sp. fig. 34a–j type locality. vestre gasværk. type stratum. middle paleocene, selandian. derivation of name. from fragilis (latin) = fragile, referring to the rather thin and fragile appearance of the otoliths of this species. holotype. fig. 34a, b, mguh 26196. paratypes. fig. 34e, f, j, vestre gasværk, selandian, mguh 26197–26199; fig. 34c, d, h, sundkrogen, selandian, mguh 26200–26201; fig. 34g, i, fakse, danian, mguh 26202–26203. diagnosis. moderately large otoliths (up to 7 mm), thin and with an oval outline. the dorsal rim is rather regularly curved, median pronounced and postdorsally reduced. the ventral rim is short with a straight medioventral portion. the anterior and posterior tips have moderately pointed angles. the sulcus is moderately wide and the ostium is slightly shorter than the cauda. further material. 208 otoliths; 86 from the danian of fakse; 54 from the selandian of sundkrogen, 82 from vestre gasværk, 3 from gemmas allé and 4 from kongedyb. measurements l h t l:h h:t 6.50 4.90 1.25 1.30 3.90 holo. 6.30 4.85 1.25 1.30 3.90 para.* 3.95 3.00 n.m. 1.30 para.* 3.90 2.85 0.80 1.35 3.55 para.† 3.75 3.05 0.85 1.25 3.60 para.* 2.85 2.15 n.m. 1.35 para. † 2.75 2.05 n.m. 1.35 para. † 1.95 1.45 n.m. 1.35 para.‡ for abbreviations used in the table, see p. 26. * from vestre gasværk. † from fakse. ‡ from sundkrogen. description. the otoliths are rounded to oval and thin and reach in size to about 7 mm. the dorsal rim is regularly curved, median pronounced and postdorsally reduced without prominent angles. the ventral rim is deeply curved, gently in juveniles, with rounded preand postdorsal angles in adults; its straight medioventral portion is rather short. preand postventral angles are located close to each other on the midventral rim. the posterior tip has a moderately pointed angle just above caudal tip. the rostrum is massive, short and blunt. the excisura and the antirostrum are feeble. the rims are delicately crenellated in juveniles and smooth in adults. the inner face is moderately convex with a long, moderately wide and slightly supramedian sulcus. the ostium is slightly shorter than the cauda; it is very wide especially ventrally and anteriorly open. the cauda is narrower than the ostium, tapering and reaching very close to the posterior tip of the otolith. the ostial colliculum is well marked and somewhat deepened; the caudal colliculum has a distinct ridge-like ventral margin. the dorsal depression is wide almost occupying the dorsal field entirely. the ventral rim is smooth and rarely with feeble indications of a ventral furrow close to the ventral rim. the outer face is slightly concave and ornamented in juveniles, but smooth in adults. all rims are sharp. ontogeny and variability. small specimens of less than 3 to 3.5 mm in length (fig. 34h–j) are intensely ornamented and more generalised in outline, whereas larger specimens (fig. 34a–g) are practically smooth. other ontogenetic changes may seem rather moderate compared to the two other berycid species described above. however, the largest specimens show the most delicate outline with the typical ventral rim and the more 79 a b c d e f g h i j fig. 34. centroberyx fragilis n. sp. a, b: holotype, selandian, vestre gasværk, mguh 26196, × 12. c–f: paratypes, selandian, vestre gasværk, mguh 26197–26199, × 12. g, i, j: paratypes, danian, fakse quarry, mguh 26201–26203, × 20. h: paratype, selandian, sundkrogen, mguh 26200, × 20. 80 pointed posterior tip. sometimes, they also exhibit a narrower cauda, which is tapering and just slightly curving towards its termination (fig. 34c, d; a very characteristic ontogenetic change observed in recent species of the genus – see schwarzhans 1981a). variability mostly concerns details of the outline and is moderate. discussion. centroberyx fragilis is easily distinguished from the simultaneously occurring hoplostethus lacinatus and also from centroberyx integer (see above). large, diagnostic and well-defined specimens show a very similar morphology in outline and habitus as the recent species. the closest relative, in my opinion, is c. subrotundus (koken 1884), which is a widespread and common species from the eocene of northern europe. their otoliths are just somewhat more compressed. differentiation of the three berycid species described from the paleocene of denmark is not always easy with small specimens (see above). order scorpaeniformes garman 1899 suborder scorpaenoidei garman 1899 family scorpaenidae risso 1827 genus scorpaena linnaeus 1758 type species. scorpaena porcus linnaeus 1758. scorpaena corallophilus n. sp. fig. 35a–f type locality. fakse quarry. type stratum. soft coral limestone, early paleocene, danian. derivation of name. referring to the association of this species with a coralline environment. holotype. fig. 35a, b, mguh 26204. paratypes. fig. 35c–f, topoand stratotype, mguh 26205–26208. diagnosis. massive and very elongated otoliths with a pointed and long rostrum and a nearly similarly pointed posterior tip. the excisura and antirostrum are well developed.the ostium is short, wide and strongly deepened. the cauda is long, narrow and swinging slightly downwards to the tapering tip, which terminates close to the postventral rim. the ventral line is sharp, but relatively short and close to the ventral rim. further material. 30 specimens from danian at fakse. measurements l h t l:h h:t 4.15 1.90 n.m. 2.20 para. 3.65 1.65 n.m. 2.20 para. 3.35 1.65 0.75 2.05 2.2 holo. 2.95 1.50 n.m. 1.95 para. 2.85 1.35 n.m. 2.10 para. for abbreviations used in the table, see p. 26. description. massive and elongated otoliths growing in sizes to somewhat over 4 mm. the ventral rim is shallow, gently curved, smooth and deepest at its middle. the dorsal rim is likewise gently and shallow curved somewhat undulating and sometimes with very indistinct midand postdorsal angles. the posterior tip is pointed. the rostrum is very long and sharp. the excisura and antirostrum are distinct. the inner face is strongly convex with a long, median and deep sulcus.the ostium is short, wide, strongly deepened and with a wide opening anteriorly. the cauda is much longer and narrower than the ostium, slightly swinging downwards with a pointed tip terminating close to the postventral rim. the dorsal depression is distinct and deep. the ventral furrow is sharp, but usually not very long and is close to the ventral rim. the outer face is slightly concave and rather smooth. the ventral rim is sharp and the dorsal rim is thick. ontogeny and variability. within the range of sizes known at present ontogenetic changes do not seem to be very pronounced. the largest otoliths, however, are also the most slender ones, but since variations in smaller ones are similar, this could also be due to certain variability. otherwise, the expression of the dorsal rim is the most variable. discussion. scorpaenid otoliths so far have rarely been recorded in the fossil record. s. corallophilus represents a very typical species of this family that can convincingly be placed in the extant genus scorpaena. roedel (1930) described another possible scorpaenid otolith as genus berycidarum marchicus from paleocene erratic boulders of frankfurt an der oder. his 81 a b c d e f gh i fig. 35. scorpaena corallophilus n. sp. a, b: holotype, danian, fakse quarry, mguh 26204, × 20. c–f: paratypes, danian, fakse quarry, mguh 26205–26208, × 20. g–i: genus scorpaenidarum marchicus (roedel 1930). holotype, selandian, erratic boulders of northern germany, pmhub, × 20. 82 a b c de f g h i j k fig. 36. a–d. genus apogonidarum sp. danian, fakse quarry, mguh 26209–26210, × 20. e–k. acropoma sp. e–g, k: selandian, vestre gasværk, mguh 26211–26212; e, k: × 20; f, g: × 12. h–j: selandian, kongedyb, mguh 26213–26215, × 20. 83 unique holotype is refigured in fig. 35g–i for comparison. genus scorpaenidarum marchicus (roedel 1930) is more compressed and shows a straight cauda. more similar is a scorpaenid species originally described as genus scorpaenidarum acutus by frost (1934) from the early eocene of england, which probably also represents the genus scorpaena. otoliths of s. acuta differ from those of s. corallophilus mainly in the shorter cauda, the less widened ostium and the deepest point of the ventral rim being anterior of the middle. order perciformes bleeker 1859 suborder percoidei bleeker 1859 family apogonidae jordan & gilbert 1882 genus indet. genus apogonidarum sp. fig. 36a–d material. 2 otoliths from the soft coral limestone, fakse quarry, danian, mguh 26209–26210. measurements l h t l:h h:t 3.10 1.75 0.70 1.75 2.5 for abbreviations used in the table, see p. 26. description. otoliths are massive, regularly oval in outline and up to about 3 mm in size. the ventral rim is gently and regularly curved; the dorsal rim is with a broadly rounded mid-dorsal angle and a less pronounced postdorsal angle. anterior and posterior rims are bluntly rounded. the rostrum is indistinct and the excisura and antirostrum are not developed. the inner face is almost flat with a rather shallow slightly inframedian sulcus. the ostium is shallow, anteriorly somewhat reduced and at its middle slightly widened. the ostial colliculum is well marked and flat. the cauda is about as long as the ostium or slightly longer, straight and with a somewhat widened and rounded tip. the caudal colliculum is somewhat deepened and distinctly separated from the ostial colliculum; it terminates at some distance from the rear tip of the cauda. the dorsal depression is distinct, wide, short and located over the collum only. the ventral field is smooth and without a ventral furrow. the outer face is convex, smooth and has a broad central umbo. discussion. these two otoliths probably represent an extinct apogonid genus with a primitive otolith morphology. characteristics for the apogonids are the lack of an excisura and antirostrum, the anteriorly reduced and dorsallywidened ostium, the straight cauda and the broad anterior to mediodorsal angle. the posterior reducedcaudal colliculum is a distinct specialised feature. these otoliths likely represent an undescribed species, but are not well enough preserved to serve as type material. family acropomatidae gill 1893 genus acropoma guenther 1859 type species. acropoma japonica guenther 1859. acropoma sp. fig. 36e–k material. 19 otoliths from the selandian; 11 from sundkrogen, 2 from vestre gasværk (mguh 26211–26212), 6 from kongedyb (figured specimens mguh 26213– 26215). measurements l h t l:h h:t ~ 4.50 2.80 0.80 3.5 3.55 2.35 n.m. 1.50 for abbreviations used in the table, see p. 26. description. the otoliths are relatively small and may grow up to 5 mm in length. they are oval with a pronounced, massive and blunt rostrum. the antirostrum and excisura are very faint or absent. the ventral rim is regularly curved and often somewhat undulating; the dorsal rim is with broad mediodorsal and more pronounced postdorsal angles; the posterior tip is blunt. the inner face is moderately convex in the horizontal direction and in the vertical direction it is almost flat and rather smooth. the sulcus is slightly supramedian, rather shallow, distinctly heterosulcoid and open to the anterior. the cauda is narrow and slightly curved towards the rounded tip; the ostium is shorter and wider with an inclined joint. the dorsal depression is rather narrow and long. the ventral furrow is always present running at some distance from the ventral rim. the outer face is flat to slightly concave and rather smooth. ontogeny. the smaller otoliths (below 3 mm) are more 84 a b c d e f g h i j k l m n fig. 37. a–e. genus carangidarum sp. selandian, sundkrogen, mguh 26216–26219, × 20. f–n. genus sparidarum sp. f–i, k–n: selandian, sundkrogen, mguh 26220–26224, × 20. j: selandian, vestre gasværk, mguh 26225, × 20. 85 rounded in outline and more generalised in appearance. specimens of more than 3 to 3.5 mm seem to be mature diagnostically. discussion. the rather flat inner face with the clear ventral furrow not very close to the ventral rim and the shape of the sulcus with its inclined ostial / caudal joint are interpreted as typical characters of the family acropomatidae, a family, not uncommon in the early tertiary fossil otolith record. these otoliths likely represent an undescribed species, but the presence of only a few and incomplete larger specimens prohibits establishing a new species. family carangidae rafinesque 1815 genus indet. genus carangidarum sp. fig. 37a–e material. 63 juvenile otoliths from the selandian, 62 from sundkrogen (figured specimens mguh 26216– 26219), 1 from kongedyb. measurements l h t l:h h:t 2.25 1.25 0.35 1.80 3.5 2.05 1.10 n.m. 1.85 2.00 1.15 n.m. 1.75 1.75 1.05 n.m. 1.65 for abbreviations used in the table, see p. 26. remarks. despite the large number of otoliths available none of them is larger then 2.5 mm. they all must be regarded as juveniles without true diagnostic features. this is also evident from the very generalised morphology of the otoliths. therefore the establishment of a new species, which it likely represents, is being postponed until more and larger specimens become available. typical for a representative of the family carangidae is the elongated shape of the thin, fragile and delicately ornamented otoliths, the narrow sulcus including the rather narrow ostium, and the downwards turned and widened caudal tip that closely approaches the postventral rim. family sparidae bonaparte 1832 genus indet. genus sparidarum sp. fig. 37f–n material. 9 otoliths; 5 from the selandian at sundkrogen (mguh 26220–26224), 2 from the selandian at vestre gasværk (mguh 26225), 2 from the danian at fakse. measurements l h t l:h h:t 3.70 2.25 0.70 1.65 3.2 * 3.20 2.10 0.55 1.55 3.8 † 2.25 1.35 0.40 1.65 3.4 * for abbreviations used in the table, see p. 26. * selandian specimen. † danian specimen. description. otoliths are up to 4–5 mm in length; they are massive and robust. the anterior and posterior tips are bluntly pointed and almost symmetrical. no antirostrum or excisura are present. the dorsal rim is rather gently curved with an indistinct mediodorsal angle, sometimes marginally crenellated. the ventral rim is more deeply, very regularly curved and smooth. the inner face is markedly convex in both directions. the sulcus is rather short, heterosulcoid and somewhat deepened. the ostium is somewhat widened and the cauda short and slightly bent downwards. the dorsal field has a faint rather small depression above the cauda; the ventral field is smooth and without a ventral furrow. the outer face is concave and rather smooth. discussion. except for the specimen of fig. 37f–h, the few other specimens available are either juveniles below 2.5 mm of length or eroded. therefore, determination of the species is postponed until more and better material comes at hand, in particular since many poorly defined fossil sparid look-like fossil otolith species have been described from the late paleocene and eocene of england, france and belgium. the status of many of those nominal species is uncertain and needs a thorough revision. this is particularly true for the species described by stinton (1965, 1966, 1978, 1980, 1984). the principal problem is that many of the early percoid otoliths morphologically are not very diversified (i.e. look very similar to each other). also, they are often small and it is not always clear whether such 86 small otoliths represent diagnostically mature specimens. in addition, documentation sometimes is too poor to allow identification without review of the original material (nolf’s and stinton’s publications). and finally, species have been established apparently almost without regard or differential diagnosis and correlation with previously described species. when reviewing previously described specimens, as for instance those of stinton (1965) from the thanetian of england, it becomes apparent that many of the species are based on inadequate juveniles and/or eroded types. this is also true for genus percidarum minimus described by roedel (1930) form the selandian of north-east germany. this species has been described based on a single eroded juvenile and in my opinion should be regarded as doubtful. suborder scombroidei bleeker 1859 family gempylidae gill 1862 genus indet. genus gempylidarum merus n. sp. fig. 38a–d type locality. fakse quarry. type stratum. soft coral limestone, danian, middle paleocene. derivation of name. merus (latin) = thin, meagre, referring to the very elongated outline of the otolith. holotype. fig. 38a, b, mguh 26226. paratype. fig. 38c, d, topoand stratotype, mguh 26227. diagnosis. very elongated and fragile otoliths with a distinct and projecting postdorsal angle behind a postdorsal concavity and an inframedian angular posterior tip. the sulcus is deep, the ostium regularly widening towards anterior, and the cauda swinging downward towards the widened and deepened tip. further material. 1 specimen from the danian, fakse. measurements l h t l:h h:t ~ 5.0 1.75 0.65 > 2.8 2.7 holo. ~ 4.7 1.90 0.50 > 2.5 3.8 para. for abbreviations used in the table, see p. 26. description. the otoliths are very elongated and fragile, particularly the thin rostrum, which is missing in all three specimens. otolith size reaches to about 5 mm or more. the ventral rim is curved very shallow and regularly. the dorsal rim is nearly flat anteriorly; thereafter it is developed as a broad concavity just before the very distinct and projecting postdorsal angle, which sits far back at the dorsal rim. the posterior rim is straight and oblique; it terminates in an inframedian angular posterior tip. the rostrum is missing, but presumably it is not very long. the antirostrum and excisura are feeble. all the rims are smooth or slightly undulating. the inner face is slightly convex with a long, rather wide and deep sulcus. the ostium is short, very deep and regularly widening and opening towards the anterior. the cauda is long, somewhat narrower than the ostium and with a widened and deepened, distinctly downwards turning tip, which terminates very close to the postventral rim. the dorsal depression is very narrow and indistinct. the ventral furrow is sometimes visible at some distance from the ventral rim. the outer face is slightly concave and smooth or with few indistinct radial furrows to the posterior side. the rims are sharp. variability. the figured paratype is somewhat thinner, less elongated and the typical posterior tip of the otolith is less well developed. discussion. scombroid otoliths are generally very rare in the fossil record. genus gempylidarum merus represents the earliest such record and it shows a very primitive otolith morphology. several typical scombroid features are weakly developed but are still recognisable. these are the distinctive shape of the posterior outline of the otolith with the projecting postdorsal angle, which sits far backwards, the shape of the ostium and the deepening and widening of the downwards turned caudal tip. in scombrids and thunnids these features are further developed and more accentuated and the ventral rim becomes nearly straight. the most primitive family in this group, the gempylidae, contains genera, which have otoliths more similar in this respect. genus gempylidarum merus probably represents an 87 a b c d e f g h i j k fig. 38. a–d. genus gempylidarum merus n. sp. a, b: holotype, danian, fakse quarry, mguh 26226, × 12. c, d: paratype, danian, fakse quarry, mguh 26227, × 12. e–h. mupus sinuosus (stinton 1965). selandian, sundkrogen, mguh 26228–26230, × 20. i–k. genus stromateoidarum sp. selandian, sundkrogen, mguh 26231, × 20. 88 extinct genus of this family, but so far not enough is known of recent gempylid otoliths to be certain. both figured otoliths lack the rostrum, but still the other characters are distinctive enough to warrant establishing of a new species. suborder stromateoidei regan 1909 family centrolophidae regan 1909 genus mupus cocco 1840 type species. mupus imperialis cocco 1840 (syn. centrolophus ovalis cuvier 1833). mupus sinuosus (stinton 1965) fig. 38e–h 1965 scombrops sinuosus stinton, p. 413, plate 32, fig. 29; plate 33, fig. 54. material. 48 otoliths (mostly broken) from the selandian; 45 from sundkrogen (figured specimens mguh 26228–26230), 3 from vestre gasværk. measurements l h t l:h h:t 2.80 1.35 0.25 2.10 5.5 for abbreviations used in the table, see p. 26. description. elongated, flat, very thin and fragile otoliths (complete specimens are in fact extremely rare). the otolith size may reach up to 5 mm. the ventral and dorsal rims are rather shallow with obtuse middorsal and mid-ventral angles. the posterior tip is rounded. the anterior tip is pointed with a long rostrum. no or very feeble excisura and antirostrum are present. the rims are undulating or crenellated. the inner face is almost flat, very slightly convex with a long, narrow and slightly supramedian sulcus. the ostium is very short in comparison to the cauda and only slightly widened. the cauda is very long, narrow, almost straight, just slightly turning downwards towards the tip, which closely approaches the posterior rim of the otolith. the dorsal depression is narrow and indistinct. the ventral field is smooth. the outer face is flat and slightly ornamented. all rims are very sharp. remarks. mupus otoliths are very distinctive and occur regularly in paleogene sediments of the north sea basin. however, they are so fragile that complete otoliths like the one shown in fig. 38h are extremely rare.m.sinuosus was originally described from the early eocene of england. other species are m. confinis nolf 1970 from the middle and late eocene of belgium and m. neumanni schwarzhans 1974 from the late oligocene of germany. m. confinis is distinguished from m. sinuosus by its somewhat more compressed outline and m. neumanni by the lack of the mid-dorsal angle. family and genus indet. genus stromateoidarum sp. fig. 38i–k material. 7 otoliths from the selandian at sundkrogen, mguh 26231. measurements l h t l:h h:t 1.85 1.50 0.30 1.25 5.0 for abbreviations used in the table, see p. 26. remarks. all specimens of this thin and fragile species are fragmented except for the single small and certainly juvenile specimen figured. this specimen is characterised by its rounded outline with its ornamented rims and the shape of the sulcus. the sulcus is almost straight including the cauda and rather narrow and somewhat deepened. the ostium is much shorter than the cauda and not much widened; the cauda reaches very close to the posterior rim of the otolith, almost opening to it in a slight concavity of the rim. the shape of the sulcus is quite typical for otoliths of the stromateoidei. the small size of the single well preserved specimen, however, does not allow for a more precise identification. order tetraodontiformes berg 1940 family ostraciidae rafinesque 1815 genus ostracion linnaeus 1758 type species. ostracion tetragonus linnaeus 1758. ostracion pergravis n. sp. fig. 39a–g 89 a b c d e g f fig. 39. ostracion pergravis n. sp. a–c: holotype, selandian, sundkrogen, mguh 26232; a: × 35; b, c: × 20. d–g: paratypes; selandian, sundkrogen, mguh 26233– 26235, × 35. type locality. sundkrogen. type stratum. selandian, paleocene. derivation of name. pergravis (latin) = very important. holotype. fig. 39a–c, mguh 26232. paratypes. fig. 39d–g, mguh 26233–26235. diagnosis. small, very high bodied, compact otoliths with nearly triangular outline, the three corners of the triangle formed by the pointed postventral, the broadly rounded preventral and the broad mid-dorsal angles. the excisura is deep and sharp; rostrum and antirostrum of about the same size, massive. the sulcus is very deep, anteriorly open and posteriorly reaching close to the posterior rim of the otolith. the ostium and cauda are about the same size, separated by a deeply intruding dorsal depression that nearly joins up with the collum. further material. sundkrogen: 3 specimens. measurements l h t l:h h:t 1.85 2.45 0.70 0.75 3.5 holo. 1.60 1.95 n.m. 0.80 para. 1.00 1.35 0.35 0.75 3.8 para. for abbreviations used in the table, see p. 26. 90 description. otoliths are rather small reaching about 2.5 mm in height, compact, thick and much higher than long with a length:height index of about 0.75. the outline is almost triangular. the ventral rim is almost straight and horizontal with a sharp postventral angle (sometimes developed into a spine); anteriorly it is curving gently upwards to the rostrum. the excisura is sharp and deep. the rostrum is very massive, short and just slightly longer than the equally massive antirostrum. the dorsal rim is very high, somewhat irregular with a prominent broad mid-dorsal angle. the posterior rim is almost straight and smooth and is somewhat inclined towards the postventral spine. all rims are smooth or slightly to irregularly undulated. the inner face is slightly convex with a very intense relief. the sulcus is very deep furrow-like and straight with a median position; anteriorly it is open and posteriorly it is reaching very close to the posterior rim of the otolith. the ostium and the cauda are about equal in size and again deepened against the rest of the sulcus, which is separated by a narrow and somewhat shallower collum. the colliculi are separated but are poorly defined. the dorsal field is occupied by a vshaped deep depression, which nearly joins up with the collum. the dorsal depression has a few radial furrows starting and radiating from a point just above the collum. the ventral rim is rather smooth except for a ventral furrow of variable intensity close to the ventral rim reaching from the rostrum to the postventral spine. the outer face is nearly flat and rather smooth except for few radial furrows on the dorsal part. the rims are thick. ontogeny and variability. ostraciid otoliths are rather small so that the holotype with its height of about 2.5 mm can be regarded as truly adult. the smallest specimen available is about half the size but still shows all the diagnostic valid features although just slightly more generalised. variability is moderate and restricted to details of the ornamentation of outline and outer side and depth of the excisura and the dorsal depression. discussion. this is the first definite fossil record of a tetraodontiform otolith and the earliest record of this family. at the same time the highly apomorphic pattern typical for otoliths of this family is already well developed. it is surprising though how close these paleocene otoliths resemble recent specimens of the ostraciid genus ostracion especially those of the subgenus lactoria. there has been one other fossil tetraodontiform otolith record in the past, i.e. amanses sulcifer stinton 1966 (family monacanthidae), but in my opinion this species rather represents a zeiform (see genus zeiformorum janni). lapilli remarks. apart from the sagittae otoliths described above the collection from the selandian of sundkrogen included very few lapillae otoliths (4 specimens) representing two different species. except for the cypriniformes and the siluriformes lapillae otoliths are much smaller than sagittae otoliths and are not regarded as specifically diagnostic by most workers. therefore, these specimens are not figured and it is not attempted to assign them to any of the species identified by sagittae otoliths. acknowledgements the otoliths from the paleocene of copenhagen and collected by harder and rosenkrantz were traced, extracted and kindly made available to me by k. ingemann schnetler, langaa and m.s. nielsen, odense. k. ingemann schnetler was also the first to arouse my interest in this extensive collection and supported me with geological, palaeontological and other information throughout my work. my very special thanks go to him for all the help. the material from the paleocene of fakse was kindly made available by mrs. a. rasmussen, fakse and k. ingemann schnetler. i also wish to thank w.-d. heinrich, berlin, who kindly made roedel’s original material from the paleocene of north-east germany available to me for revision. a further collection of paleocene otoliths from west greenland was traced and made available by k. ingemann schnetler from the collection of a. rosenkrantz catalogued at the geological museum of the university of copenhagen (mguh). these otoliths are the subject of a separate study currently under preparation. finally, i wish to thank d. nolf, brussels and mrs. b. reichenbacher, karlsruhe for the critical review of the manuscript and many useful recommendations. 91 references agassiz, l.r. 1843: recherches sur les poissons fossiles, 90 pp. neuchâtel. andersen, s.b. & heilmann-clausen, c. 1984: petrografi og alder af den brune turitella-sandsten, en tertiær løsblok fra østersøområdet. dansk geologisk forening, årsskrift for 1983, 17– 24 (with english abstract). barnard, k.h. 1934: new records and descriptions of two new species of south african marine fishes. annual magazine natural history 13, 228–235. berg, l.s. 1940: sistema ryb. trudy zoologicheskovo instituta akademii, 517 pp. moscow: nauk sssr. berggren, w.a. 1994: in defense of the selandian 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geological atlas of western and central europe, 2nd edition, 239 pp. the hague: shell internationale petroleum maatschappij, b.v. 95 danmarks og grønlands geologiske undersøgelse (geus) geological survey of denmark and greenland øster voldgade 10, dk-1350 copenhagen k denmark geological survey of denmark and greenland bulletin is a new series started in 2003 to replace the two former bulletin series of the survey, viz. geology of greenland survey bulletin and geology of denmark survey bulletin. the twenty-one volumes published since 1997 in those two series are listed below, followed by titles in the new bulletin series. the new series, together with geological survey of denmark and greenland map series, now form the peer-review scientific series of the survey. geology of greenland survey bulletin (discontinued) 173 cambrian shelf stratigraphy of north greenland, 120 pp., 1997. by j.r. ineson & j.s. peel. 250.00 174 the proterozoic thule supergroup, greenland and canada: history, lithostratigraphy and development, 150 pp., 1997. by p.r. dawes. 300.00 175 stratigraphy of the neill klinter group; a lower – lower middle jurassic tidal embayment succession, jameson land, east greenland, 80 pp., 1998. by g. dam & f. surlyk. 250.00 176 review of greenland activities 1996, 112 pp. (18 articles), 1997. edited by a.k. higgins & j.r. ineson. 200.00 177 accretion and evolution of an archaean high-grade grey gneiss – amphibolite complex: the fiskefjord area, southern west greenland, 115 pp., 1997. by a.a. garde. 200.00 178 lithostratigraphy, sedimentary evolution and sequence stratigraphy of the upper proterozoic lyell land group (eleonore bay supergroup) of east and north-east greenland, 60 pp., 1997. by h. tirsgaard & m. sønderholm. 200.00 179 the citronen fjord massive sulphide deposit, peary land, north greenland: discovery, stratigraphy, mineralization and structural setting, 40 pp., 1998. by f.w. van der stijl & g.z. mosher. 200.00 180 review of greenland activities 1997, 176 pp. (26 articles), 1998. edited by a.k. higgins & w.s. watt. 200.00 181 precambrian geology of the disko bugt region, west greenland, 179 pp. (15 articles), 1999. edited by f. kalsbeek. 240.00 182 vertebrate remains from upper silurian – lower devonian beds of hall land, north greenland, 80 pp., 1999. by h. blom. 120.00 183 review of greenland activities 1998, 81 pp. (10 articles), 1999. edited by a.k. higgins & w.s. watt. 200.00 184 collected research papers: palaeontology, geochronology, geochemistry, 62 pp. (6 articles), 1999. 150.00 185 greenland from archaean to quaternary. descriptive text to the geological map of greenland, 1:2 500 000, 93 pp., 2000. by n. henriksen, a.k. higgins, f. kalsbeek & t.c.r. pulvertaft. 225.00 186 review of greenland activities 1999, 105 pp. (13 articles), 2000. edited by p.r. dawes & a.k. higgins. 225.00 187 palynology and deposition in the wandel sea basin, eastern north greenland, 101 pp. (6 articles), 2000. edited by l. stemmerik. 160.00 188 the structure of the cretaceous–palaeogene sedimentary-volcanic area of svartenhuk halvø, central west greenland, 40 pp., 2000. by j. gutzon larsen & t.c.r. pulvertaft. 130.00 189 review of greenland activities 2000, 131 pp. (17 articles), 2001. edited by a.k. higgins & k. secher. 160.00 190 the ilímaussaq alkaline complex, south greenland: status of mineralogical research with new results, 167 pp. (19 articles), 2001. edited by h. sørensen. 160.00 191 review of greenland activities 2001, 161 pp. (20 articles), 2002. edited by a.k. higgins, k. secher & m. sønderholm. 200.00 96 geology of denmark survey bulletin (discontinued) 36 petroleum potential and depositional environments of middle jurassic coals and non-marine deposits, danish central graben, with special reference to the søgne basin, 78 pp., 1998. by h.i. petersen, j. andsbjerg, j.a. bojesen-koefoed, h.p. nytoft & p. rosenberg. 250.00 37 the selandian (paleocene) mollusc fauna from copenhagen, denmark: the poul harder 1920 collection, 85 pp., 2001. by k.i. schnetler. 150.00 geological survey of denmark and greenland bulletin (new series) 1 the jurassic of denmark and greenland, 948 pp. 2003. edited by j.r. ineson & f. surlyk. 2 fish otoliths from the paleocene of denmark, 94 pp. 2003. by w. schwarzhans. forthcoming volumes late quaternary environmental changes recorded in the danish marine molluscan faunas. by k.s. pedersen. the jurassic of north-east greenland. edited by l. stemmerik & s. stouge. review of survey activities, 2003. edited by a.k. higgins & m. sønderholm. the lopra deepening project, faeroe islands. edited by r. waagstein & s. stouge. prices are in danish kroner exclusive of local taxes, postage and handling geological survey of denmark and greenland bulletin 4, 2003, pp 9-12 9 in an oil reservoir, the geometry of the interface between water and oil is critical in determining the volume of oil trapped below the top seal. if the interface is planar and horizontal, the volume calculation is fairly simple, but if the interface is tilted or undulating, estimation of the volume of the trapped oil is complex as it depends on the combined structural and fluid contact geometry. since accumulation of the oil may take place over a time span of several million years, while the reservoir is experiencing burial and compaction, the charge history must be studied using dynamic methods that account for these changes and for flow in both the oil and water phases. these processes have been studied quantitatively at the geological survey of denmark and greenland (geus) in a project that has combined the burial model with a fluid flow simulator. the modelling study shows that filling of a chalk reservoir can have a very long and complex history dominated by very low fluid flow rates (cm/year). the resulting modelled present-day situation exhibits a very irregular oil distribution and a non-planar geometry of the fluid contacts, and shows marked similarities to that shown by the field data. oil–water contact and free water level the positions of the oil–water contact (owc), the gas–oil contact (goc) and the associated free water level (fwl) in an oiland gas-field are some of the most important factors in estimating the in-place hydrocarbon volumes of a given field. thus it is important to be able to analyse and predict tilted or irregular fluid contacts (dennis et al. 2000; moss et al. 2003; dennis et al. in press; vejbæk et al. in press). the fluid contact can be defined in two radically different ways: the owc is defined by setting a threshold for the oil saturation, whereas the fwl is defined where the pressures the history of hydrocarbon filling of danish chalk fields peter frykman, ole v. vejbæk, niels bech and carsten m. nielsen fig. 1. map showing top chalk depth structure for the danish north sea area. producing chalk fields are shown, with oil fields green, and gas fields red. colour interval is 100 m and contour interval 50 m. the red line on the kraka field shows location of the profile studied (see fig. 3). full black lines are major faults. dashed black lines are offshore sector boundaries. modified from vejbæk et al. (in press). geological survey of denmark and greenland bulletin 4, 9–12 (2004) © geus, 2004 10 in the water and the oil phases are equal. in the chalk reservoirs in the north sea, the relationship between the owc and the fwl can be described in simple cases by the capillary characteristics of the reservoir rock. in the central north sea (fig. 1), the fluid contacts in the chalk can be naturally tilted by hydrodynamic activity due to a regional flow of water in the chalk. a regional pressure gradient in the chalk aquifer has been described from available pressure measurements (megson 1992), and later refined using more data (dennis et al. in press). the regional lateral pressure gradient reflects differential compaction caused by rapid neogene deposition with the highest burial rates in the central ekofisk area (japsen 1998). the water therefore migrates laterally away from this area and towards the periphery of the north sea. analysis of burial history by backstripping and decompaction shows that this pressure was probably caused mainly by rapid deposition in the time interval from latest miocene to recent times, as the magnitude of the pressure corresponds to the thickness of these deposits (japsen 1998). this is consistent with a very low regional permeability of the chalk (and adjacent sedimentary packages) probably not exceeding 1 md. the flow of water and the accompanying pressure differences will influence the position of the fwl (fig. 2a). if the oil is also flowing due to either buoyancy equilibration or active migration, it will affect both the fwl and the goc (fig. 2b, c). factors that modify the position of the fwl include tilting due to structural movements, and the presence of oil migrating from the underlying source rocks into the reservoir. the reason that these processes influence the present geometry of the fwl is that both oil and water flow take place at very low velocities (cm/year), due to the low permeability of the chalk. even though structural movements are very slow, the flow is not able to respond quickly enough to equilibrate the system, even on a scale of millions of years. the petrophysical properties of the north sea chalk reservoirs are mainly governed by their high-porosity/low-permeability aspect with porosities usually around 20–40% and average permeability of 1 md. case study the kraka field in the southern danish north sea (fig. 1) has been chosen as the subject of a case study of primary oil charging and remigration. to study the interaction of the different processes, reservoir fluid flow simulation techniques have been applied in combination with burial modelling, including compaction (vejbæk 2002). the results show that a time span in the order of 2 ma is required for the hydrocarbons to reach the top of the reservoir in an approximately equilibrium state, if they enter the reservoir section from a flank position. however, not even dynamic equilibrium can be fully obtained in this time span if re-perturbation by structural movements leads to changing water-zone pressure gradients. the study is focused on a 2d section from the crestal part through the south-eastern flank of the kraka field (figs 1, 3). since porosity is the main cause for changes in seismic impedance (japsen et al. in press), detailed porosity profiles can be achieved by converting acoustic impedance derived by seismic inversion. these porosity profiles have been modified by backstripping to reconstruct geometry and porosity. as flow simulation has only been applied to the chalk group layers, fig. 2. possible dynamic equilibrium situations that may fit a tilted oil–water contact. arrows show direction of pressure drop corresponding to flow direction. 1 and 2 represent wells where the pressure depth plots shown to the left are generated: (a) only the water phase is dynamic; (b) both oil and water are flowing, but the tilt is maintained due to a higher lateral pressure gradient in the water phase; (c) the tilt is maintained only by an oil phase gradient. the situations are physically distinguishable by the dip of the gas–oil contact. modified from vejbæk et al. (in press). 11 detailed porosity profiles have only been constructed for these layers. the simulation of flow processes in chalk reservoirs is characterised by the need for end-point scaling and hysteresis, in order to account for the marked influence from the high capillary forces in this low-permeability medium. since the dominant process during the filling history is a drainage process (i.e. oil replacing water), the saturation functions must also be derived for this type of process. there is a general lack of relative permeability analyses for drainage, and therefore imbibition curves have been the guide for establishing the drainage saturation functions for relative permeability. for each rock type (danian and maastrichtian), the irreducible water saturation (swi) and capillary entry pressure (pce) are assumed to depend upon the porosity (φ) through relatively simple relationships. using these relationships, the primary drainage capillary pressure is described by means of the eqr model (engstrøm 1995). the simulation of the filling history uses 8 million years before present as the starting point, and the entry of hydrocarbons from an underlying source rock is assumed to occur on the south-eastern flank (fig. 3). the flow simulation of the filling dynamics of the kraka chalk reservoir has a complex geometry due to the high capillary entry pressures in the low-permeability chalks. these internal barriers re-direct hydrocarbons, such that oil flows in the maastrichtian layers for some time before it is able to penetrate upwards into the overlying danian chalk (fig. 3a). if oil supply is stopped after 1 million years, the oil continues to move towards the crest, but leaves immobile residual oil on the migration route. hydrocarbon charging is slow and equilibration of hydrocarbons with respect to pressure gradients therefore occurs very slowly. after two million years, the oil is seen to be nearly in equilibrium even though the fwls are still slightly inclined and do not coincide for the two reservoir units (fig. 3b). after 4 million years, equilibrium is more obvious (fig. 3c). between 2 million years before present and the present, a fig. 3. modelled oil saturation in the kraka field profile at different times during the simulated filling history. top structures of the ekofisk (maastrichtian) and tor (danian) formations are shown as thin green and red lines, respectively. calculated free water levels (fwls) for these two reservoir units are shown in thicker green and red lines, respectively. charging of the reservoir starts at 8 ma b.p. by injecting oil at a very low rate at the flank position shown with an arrow. a: the situation after 250 000 years, where the injected oil is preferentially moving in the maastrichtian reservoir unit. b: oil distribution after 2 ma, where the charging has been sustained over the first 1 ma, accompanied by equilibration during continued burial. c: oil is near-equilibrium at 4 ma b.p. d: tilted fwls resulting from a lateral pressure gradient of 3.5 psi/km (24.1 kpa/km) applied for 1 ma (from 2 to 1 ma b.p.) within the water phase. this causes a water-flow south-eastwards in the aquifer, and accordingly a tilting of the fwl in that direction, which is further accentuated until the modelled present-day situation shown in e. modified from vejbæk et al. (in press). 12 pressure gradient is imposed in the water zone in order to allow for the regional pressure distribution during that period. as a result, the oil is forced south-eastwards towards the flank (fig. 3d), which is further accentuated through time as the water gradient is sustained (fig. 3e). again a zone with residual oil is left behind. conclusions the modelling reported here demonstrates that oil accumulations in chalk may require several million years to equilibrate following perturbations resulting from primary migration or reservoir tilting, if matrix permeability governs fluid flow. since naturally occurring disequilibrium oil accumulations dominate the danish chalk fields, it must be concluded that matrix flow dominates fluid dynamics. the modelled filling scenarios are intended to illustrate the general aspects of geological timescale oil–water dynamics in chalk reservoirs. the scenarios are not considered to represent actual filling histories, as they are constrained by relatively simple model assumptions, but they are geologically plausible. due to the long equilibration times, it can be dangerous to interpret tilted contacts as reflecting only dynamic equilibrium, as they may be fully dynamic and still actively flowing. this is revealed locally by non-equilibrium between danian and maastrichtian oil where they are seen to have different fwls. it is important to try to understand fluid dynamics during exploration work, since this strongly affects trap definition and volumes. the project shows that with simple and geologically based assumptions, a reasonable filling history can be modelled quantitatively. a reasonable end-result can be produced that has many similarities with present-day hydrocarbon configurations. with the methods developed in the project, even a fully dynamic system (with both oil and water moving), as for example in the dan–halfdan field system, may be explained. acknowledgement the work presented in this paper was partly funded by the danish energy authority (grant no. 1313/01-0004). references dennis, h., baillie, j., holt, t. & wessel-berg, d. 2000: hydrodynamic activity and tilted oil–water contacts in the north sea. in: ofstad, k., kittilsen, e.-j. & alexander-marrack, p. (eds): improving the exploration process by learning from the past. norwegian petroleum society (npf), special publications 9, 171–185. dennis, h., bergmo, p. & holt, t. in press: tilted oil–water contacts – modelling the effects of aquifer heterogeneity. in: doré a.g. & vining, b. (eds): petroleum geology: north-west europe and global perspectives. proceedings of the 6th petroleum geology conference. london: geological society. engstrøm, f. 1995: a new method to normalize capillary pressure curves. 1995 international symposium of the society of core analysts, san francisco, ca, usa, september 12–14, 1995, sca-9535, 12 pp. japsen, p. 1998: regional velocity-depth anomalies, north sea chalk; a record of overpressure and neogene uplift and erosion. american association of petroleum geologists bulletin 82(11), 2031–2074. japsen, p., bruun, a., fabricius, i.l., rasmussen, r., vejbæk, o.v., pedersen, j.m., mavko, g. & mogensen, c. in press: influence of porosity and pore fluid on acoustic properties of chalk: avo-response from oil, south-arne field, north sea. petroleum geoscience. megson, j.b. 1992: the north sea chalk play: examples from the danish central graben. in: hardman, r.f.p. (ed.): exploration britain: geological insights for the next decade. geological society (london) special publication 67, 247–282. moss, b., barson, d., rakhit, k., dennis, h. & swarbrick, r. 2003: formation pore pressures and formation waters. in: evans, d. et al. (eds): the millenium atlas: petroleum geology of the central and northern north sea, 317–329. london: geological society. vejbæk, o.v. 2002: a 1, 2 and 3d backstripping procedure with application to the kraka field. danmarks og grønlands geologiske undersøgelse rapport 2002/44, 35 pp. vejbæk, o.v., frykman, p., bech, n. & nielsen, c.m. in press: the history of hydrocarbon filling of chalk fields. in: doré, a.g. & vining, b. (eds): petroleum geology: north-west europe and global perspectives. proceedings of the 6th petroleum geology conference. london: geological society. authors’ addresses p.f., o.v.v. & n.b., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: pfr@geus.dk c.m.n., danish energy authority, amaliegade 44, dk-1256 copenhagen k, denmark. << /ascii85encodepages false /allowtransparency false /autopositionepsfiles true /autorotatepages /all /binding /left /calgrayprofile (dot gain 20%) /calrgbprofile (srgb iec61966-2.1) /calcmykprofile (u.s. web coated \050swop\051 v2) /srgbprofile (srgb iec61966-2.1) /cannotembedfontpolicy /warning /compatibilitylevel 1.4 /compressobjects /tags /compresspages true /convertimagestoindexed true /passthroughjpegimages true /createjdffile false /createjobticket false /defaultrenderingintent /default /detectblends true /colorconversionstrategy /leavecolorunchanged /dothumbnails false /embedallfonts true 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<feff4f7f752890194e9b8a2d5b9a5efa7acb76840020005000440046002065874ef65305542b8f039ad876845f7150cf89e367905ea6ff0c4fbf65bc63d066075217537054c18cea3002005000440046002065874ef653ef4ee54f7f75280020004100630072006f0062006100740020548c002000520065006100640065007200200035002e0030002053ca66f465b07248672c4f86958b555f3002> >> >> setdistillerparams << /hwresolution [2400 2400] /pagesize [595.000 842.000] >> setpagedevice geological survey of denmark and greenland bulletin 31, 2014, 27-30 27 a novel technique for obtaining representative water samples during co2 core-flooding experiments on chalk at reservoir conditions claus kjøller and john zuta there is a huge potential for using co2 gas to recover additional oil after water flooding in reservoir chalk. however, the injection of co2 into chalk reservoirs will disturb the chemical equilibrium between formation water, injection water and chalk. a proper understanding of these co2induced interactions and the resulting changes in the physical properties at representative reservoir conditions is required. unfortunately, reliable chemical data are rare because data cannot be acquired directly at reservoir conditions with present-day techniques. in published experiments, water samples are in many cases obtained at atmospheric conditions with the aid of a back-pressure regulator. thus, water samples are not representative of in situ reservoir conditions and if proper care is not taken, the collected data cannot be used to judge the magnitude of the chemical reactions taking place at reservoir conditions. however, in some cases water obtained at laboratory conditions can give information on in situ reservoir conditions by using geochemical speciation models to account for dissolved gases that are lost from the effluents during sampling (bachu & adams 2003). the objective of this study was to develop and test a new technique for obtaining water samples during co2-brine– rock interactions in reservoir chalk under representative reservoir conditions and gain a better understanding of the chemical interactions which occur during the injection of co2. the experiments were performed by injecting co2-saturated seawater at supercritical co2 conditions, at a pressure of 17.24 mpa (2500 psig; pounds per square inch of gauge pressure) and a temperature of 75°c. these values are typical of a region in a chalk field where the rock is exposed to long term reactions with flow of co2-bearing water. in addition, the numerical code phreeqc-3 (parkhurst & appelo 2013) was used to reproduce the experiments by assuming equilibrium between calcite and injected fluids. in this way, we can validate the sampling procedure and investigate how the measured parameters (ph, co2 pressure, calcium (ca) and bicarbonate (hco3 –) concentrations) compared with the calculated parameters. reservoir chalk samples chalk plugs for the experiment were sampled from the maastrichtian tor formation in the central north sea and were first cleaned of salt and oil with methanol and toluene. following initial determination of porosity and permeability, the samples were saturated with synthetic formation water (fw). all plugs had a diameter of 3.81 cm and a length of approximately 7.5 cm; plugs with similar petrophysical properties (table 1) were used to check the reproducibility of the studied sampling procedure. the compositions of synthetic formation water, synthetic seawater and co2 synthetic saturated seawater are listed in table 2. experimental setup the setup/rig for the experiments is shown in fig. 1. it includes three cylinders with pistons placed in an oven to maintain constant temperature. two of the cylinders were used © 2014 geus. geological survey of denmark and greenland bulletin 31, 27–30. open access: www.geus.dk/publications/bull waste co2-saturated waterlimit of oven seawater flooding plug eq ui l. p lu g primary sample loop (13 ml) separation point secondary sample cylinder (125 ml) pressure transducer pressure transducer ph-meter pressure transducer st1 st2 pt1pt2 primary sample loop (13 ml) fig. 1. experimental setup/rig within the oven. the rig has two primary sample loops. 2828 for the delivery of co2-saturated seawater and seawater. the third cylinder was used as a waste tank and for maintaining fluid pressure. all the fluids were injected at a constant rate of 6.5 ml/h with the plug in a horizontal position. the injection of co2-saturated seawater was preceded by injecting close to 3.5 pore volume (pv) of seawater to mimic the present-day water flooding. at the end of the injection of co2-saturated seawater, seawater was again injected through the plug. at this point, equilibrium with calcium carbonate was obtained in the injection fluid by first passing the seawater through the equilibrium plug (fig. 1). this was done in order to avoid any further dissolution of chalk in the plug under study. the final seawater flooding was performed to displace any residual co2 gas and bring the concentrations back to initial levels. water sampling technique the system for obtaining water samples was located at the downstream end of the rig (fig. 1). it consists of two primary sample loops each with a volume of c. 13 ml. this enabled continuous collection of water samples for every 13 ml (c. 0.5 pore volume) of plug flooded. the sampling was done by alternating the flow process between the two primary loops with the aid of computer-controlled valves (cv-210 valves) without disrupting the injection process. the primary sample loops were connected to a secondary sample cylinder with a volume of 125 ml by a separation point. water samples were transferred from the primary sample loop to the secondary sample cylinder by a de-pressurisation step. this made it possible to collect water samples for a set of chemical parameters (ca and hco3 – concentrations) at predefined sampling intervals. the ph and co2 pressures of the de-pressurised effluents were measured on-line in the secondary sample cylinder after the de-pressurisation step. the sampling from the secondary sample cylinder was subsequently done as fast as possible, in less than 5 min., thereby providing a consistent chemical dataset for the conditions prevailing in the secondary sample cylinder. in this way it was possible to validate the subsequent numerical calculations used to estimate the in situ chemical conditions. a detailed description of the sampling procedure as well as documentation for the consistency of the collected dataset was published by kjøller & zuta (2012). batch modelling with phreeqc-3 the modelling was performed at the same conditions – pressure and temperature of 17.24 mpa (2500 psig) and 75°c as the experiments, in order to estimate the true in situ reservoir condition chemistry. three equilibrium steps in a closed system were investigated: (1) equilibrium between co2 and seawater corresponding to the injected solution, (2) equilibrium between co2-saturated seawater and calcite, representing reservoir conditions, and (3) composition of the solution after de-pressurisation in the secondary sample cylinder. a good fit between experimental and numerical data after step table 1. petrophyscial properties of plugs prior to the injection of co2-saturated water at 17.24 mpa (2500 psig) and 75°c 17a 32.60 26.94 1.36 19a 30.60 25.59 1.49 porosity pore volume permeability plug id (%) (ml) (md) table 2. composition (in mg/l) of synthetic formation water (fw), seawater (sw), and co2-saturated seawater (csw) na 22866 11090 11090 k 175 408 408 mg 226 1370 1370 ca 1244 434 434 sr 142 6.8 6.8 cl– 38383 20173 20173 hco3 – 22 30 70 so4 2– – 2780 2780 ions (fw) (sw) (csw) 0 0.033 0.067 0.100 0.133 0.167 0.200 0.233 4 4.5 5 5.5 6 6.5 7 7.5 0 5 10 15 20 25 c o 2 p re ss ur e (m pa ) ph pore volume ph plug 17a ph plug 19a co2 pressure plug 17a co2 pressure plug 19a fig. 2. measured ph and co2 pressure for plug 17a and parallel plug 19a versus cumulative, injected pore volume. ph 3.2 4.8 6.3 ca (mg/l) 438 1970 1970 alkalinity, hco3 – (mg/l) 82 3274 3308 saturation index (calcite) –3.78 0 1.43 chemical parameter step 1 step 2 step 3 table 3. composition of resulting solutions at the two different equilibrium steps followed by the de-pressurisation step based on calculations with phreeqc.dat database in phreeqc-3 29 3 is considered as a validation of the numerical estimate of reservoir conditions in step 2 (kjøller & zuta 2012). results figures 2–4 show ph, co2 pressure, ca, and hco3 – concentrations measured as a function of injected cumulative pore volumes. the compositions of the resulting solutions calculated at the three steps with phreeqc-3 are shown in table 3. equilibrium of co2 and seawater results in an initial co2-saturated seawater solution with a ph of 3.2 and a saturation index of calcite of –3.78 (table 3, step 1), an indication of a strong dissolution potential for chalk. thus, co2 is dissolved and reacts with water to form carbonic acid (h2co3). it is part of this acidity that is removed as the dissolution process advances into the plugs. the dissolution of calcium carbonate in the plugs takes place according to the overall chemical reaction: during the injection of seawater within the first 3.5 pore volume, the ph increases to between 7.5 and 7.7 and subsequently decreases to an average of 6.1 during the injection of co2-saturated seawater (fig. 2). at the final stage of the experiment, the ph increases again to the initial seawater ph level, an indication that all the residual co2 was removed from the plugs during the injection of seawater in equilibrium with chalk. the calculated decrease in ph at reservoir conditions (table 3, step 2) corresponds with the measured decrease in ph after 3.5 pore volume, where breakthrough of the co2-saturated seawater is observed. however, the calculated reservoir condition ph of 4.8 is much lower than the average measured ph during the injection of co2-saturated seawater. this was expected since the ph was measured after the de-pressurisation step where co2 gas had already degassed from the solution. thus, taking the ph (6.3) of the degassed solution into account in the numerical simulation provides an excellent fit to the average measured ph of 6.1 (table 3, step 3). the co2 pressure in the secondary sample cylinder varied between 0.17 mpa (25 psig) and 0.24 mpa (35 psig) at the de-pressurisation stage during the injection of co2-saturated seawater (fig. 2). this matches with the calculated co2 pressure of 0.20 mpa (28.9 psig) in the gas phase of the secondary sample cylinder (table 4), and further validates the numerical calculations. according to the equation, degassing of co2 from the water samples, caused by the transfer of water samples from the primary sample loop to the secondary sample cylinder, was expected to result in precipitation of caco3. comparison of ca and hco3 – concentrations measured in both filtered and unfiltered water samples showed no evidence of mobilisation or precipitation of fine particles during the de-pressurisation stage, neither in plug 17a nor in plug 19a (fig. 3a, b). the ca concentration decreases during the first 3.5 pore volume to the ca-concentration level in seawater h co caco ca 2hco2 3 3 2 3+ +" + table 4. properties of the gas phase generated after the de-pressurisation stage in the 125 ml secondary sample cylinder with the phreeqc.dat database in phreeqc-3 total pressure 2.00 mpa (28.9 psig) gas volume 1.25e-001 litres molar volume 1.38e+001 litres/mole p×vm/rt 0.99228 (compressibility factor z) 0 500 1000 1500 2000 2500 3000 3500 0 500 1000 1500 2000 2500 0 5 10 15 20 25 h c o 3– co nc en tr at io n (m g/ l) c a co nc en tr at io n (m g/ l) pore volume ca filt ca ca sw ca fw hco3 – filt hco3 – a fig. 3. measured ca and hco3 – concentrations for filtered and unfiltered water samples in plugs 17a (a) and 19a (b) versus cumulative, injected pore volume. filt: filtered. sw: synthetic seawater. fw: synthetic formation water. 0 500 1000 1500 2000 2500 3000 3500 0 500 1000 1500 2000 2500 0 5 10 15 20 25 c a co nc en tr at io n (m g/ l) ca filt ca ca sw ca fw hco3 – filt hco3 – h c o 3– co nc en tr at io n (m g/ l) pore volume b 3030 of c. 450 mg/l. subsequently, the measured ca concentration increases to an average of 1200–1300 mg/l during the injection of co2-saturated seawater, which is due to dissolution of calcium carbonate in the plugs. at the final steps of the experiment, the ca concentration again decreases to the seawater level during injection of seawater to remove residual co2 from the plugs. in the geochemical model, the calculated ca concentration of 1970 mg/l at reservoir conditions (table 3, step 2) is simulated by dissolution of calcite at reservoir conditions, and does not take into account the possible precipitation of calcium carbonate during sampling. therefore, the calculated reservoir-condition ca concentration is higher than the concentrations measured in the experiments during the period with co2-saturated seawater flooding. since no mobilisation or deposition of suspended fine particles in the water samples were observed, the difference in ca concentration must be attributed to deposition of fine calcium carbonate particles on the inner surface walls of the secondary sample cylinder during the pressure-reduction stages. in line with this, the simulated saturation index (si) with respect to calcite of 1.43 after de-pressurisation (table 3, step 3) is higher than the saturation indices between 0.6 and 1.2 that can be calculated based on the measured water chemistry. if the saturation index with respect to calcite is constrained to be between 0.6 and 1.2 in the numerical calculations in step 3, the calculated ca concentration at the sampling conditions varies between c.1150–1650 mg/l, which is in much better agreement with the actually measured ca concentrations (fig. 3). similar considerations and conclusions can be made for the hco3 – -concentration trend during the experiment. thus, the measured hco3 – concentration averages 2350–2550 mg/l during the injection of co2 saturated seawater, while the calculated hco3 – concentration after de-pressurisation is 3308 mg/l when no constraints on the saturation index with respect to calcite are applied (table 3, step 3). however, accounting for some calcium carbonate precipitation during de-pressurisation will diminish the difference between calculated and measured concentrations, and overall it is suggested that the calculated reservoir condition concentrations shown in table 3, step 2 represent the true in situ reservoir conditions prevailing in the plugs during experiments. the general decrease in measured ca and hco3 – concentrations during the injection of co2-saturated seawater from 3.5 to 15 pore volume (fig. 3), suggests that, with time, there is an increasing amount of calcium carbonate precipitating in the secondary sample cylinder during sampling. for future studies, this gradual change in calcium carbonate precipitation should be avoided in order to minimise the uncertainty in the measured data, and thereby also the uncertainty of the numerical estimate of the reservoir condition chemistry (table 3, step 2). conclusions • a new sampling technique was developed and used to study co2-brine–rock interactions in reservoir chalk at reservoir conditions. • the similar results obtained in the parallel plugs validate the sampling procedure and show a high degree of reproducibility. • the results from the experiment and the modeling show increased levels of calcium and alkalinity (hco3 –) with decreasing ph during the injection of co2-saturated water at reservoir conditions. there were, however, differences between the measured and calculated calcium concentration and hco3 – concentration which is probably due to accumulation of fine calcium carbonate particles on the inner walls of the secondary sample cylinder during the de-pressurisation stage. • further refinement of the method may include modification of the secondary sample cylinder, in order to minimise the accumulation of fine calcium carbonate particles on its inner surface walls. acknowledgements we are grateful to dansk undergrunds consortium (duc) – a co-operation between a.p. møller maersk, shell, chevron and nordsøfonden (the danish north sea fund) – for sponsoring the work. references bachu, s. & adams, j.j. 2003: sequestration of co2 in geological media in response to climate change: capacity of deep saline aquifers to sequester co2 in solution. energy conversion and management 44, 3151–3175. kjøller, c. & zuta, j. 2012: co2-brine–rock interactions in reservoir chalk rock – a coupled experimental and numerical approach for obtaining hydrochemical results at reservoir conditions. paper sca201204 presented at international symposium of the society of core analysts, aberdeen, scotland, uk, 27–30 august, 2012, 12 pp. parkhurst, d.l. & appelo, c.a.j. 2013: description of input and examples for phreeqc version 3 – a computer program for speciation, batch-reaction, one-dimensional transport, and inverse geochemical calculations. u.s. geological survey techniques and methods, http:// pubs.usgs.gov/tm/06/a43/. authors’ address: geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: clkj@geus.dk geological survey of denmark and greenland bulletin 31, 2014, 19-22 19 the lower palaeozoic shale gas play in denmark niels h. schovsbo, arne t. nielsen and donald l. gautier the unconventional gas resources in the lower palaeozoic shale of denmark were recently assessed by the united states geological survey (usgs; gautier et al. 2013). assuming unrestricted application of best practice current technology, recoverable gas resources of 0 to 130 × 109 nm3 gas were estimated onshore (mean = 67 × 109 nm3 gas) and 0 to 228 × 109 nm3 gas were estimated offshore (mean = 119 × 109 nm3 gas), i.e. a total estimated mean of 186 × 109 nm3 gas (nm3: normal cubic metre, unit used for natural gas at 0°c and 101.325 kpa). nearly all of this potential resource is assumed to be contained in the cambro-ordovician alum shale. the wide range of estimates reflects the sparse data and the geological uncertainty inherent in the still untested play. the estimated mean quantity of gas resource is comparable to the total volume of gas produced from the danish part of the north sea during 1972–2011 and twice the amount of the estimated remaining reserves of conventional gas in the danish part of the north sea. the assessment is the result of collaboration between the geological survey of denmark and greenland (geus) and usgs. geus and the university of copenhagen contributed with the geological input data and models and usgs provided assessment methodology and north american resource analogues. in this paper the geological model that underlies the assessment is presented along with some additional considerations on the nature of the play. details and methodology of the assessment itself were summarised by gautier et al. (2013). the danish shale gas play shale gas is an unconventional energy resource in which gas is produced directly from a shale source rock. highly productive formations in north america are regionally extensive, tens of metres thick, highly organic-rich, and have been buried sufficiently to reach the temperatures necessary for thermal gas generation. the technique of using horizontal drilling and hydraulic fracturing to extract gas from shale was developed in north america, where it has had a significant impact on gas markets. application of this technology has not yet led to shale gas production in europe. the first shale gas exploration borehole in denmark, the vendsyssel-1, is due to be drilled in northern jylland in 2015 by the company total e&p (fig. 1). since 2009, geus has conducted a wide range of shale gas evaluation programmes including screening of onshore denmark for potential shale gas units. the evaluation is partly based on extensive shallow coring on bornholm where the shale is accessible immediately beneath a thin quaternary cover (schovsbo et al. 2011). the main target for exploration in denmark is the alum shale formation, which is up to 180 m thick and unusually rich in organic matter, typically with 5–10% total organic carbon (toc; schovsbo et al. 2011). organic-rich shales also occur in younger ordovician– silurian successions. these black shales are thinner and less toc-rich than the alum shale, but may still be interesting for shale gas exploration. thermal modelling the terne-1 borehole, drilled in kattegat in 1985 (fig. 1), penetrated a 180 m thick alum shale formation and is a © 2014 geus. geological survey of denmark and greenland bulletin 31, 19–22. open access: www.geus.dk/publications/bull fig. 1. map showing simplified distribution of lower palaeozoic strata in denmark and the location of scientific and exploration boreholes used for the geological assessment. the position of the planned vendsyssel-1 borehole is also shown. in all boreholes made so far in denmark and skåne, the alum shale is mature to gas rank and positioned beneath a palaeozoic sequence less than 1 km thick; at these sites the shale did not contain significant amounts of gas. the distribution of lower palaeozoic strata is from nielsen & schovsbo (2011). 100 km bornholm skåne sweden germany kattegat c4 a3 b2 terne-1norwegian–danish basinringkøbing–fyn high jylland lower palaeozoic strata caledonian front borehole sæby-1 vendsyssel-1 slagelse-1 2020 key borehole for shale gas exploration in denmark. thermal modelling of terne-1 was carried out to calibrate the burial history and maturation profiles. the alum shale contains a marine type ii kerogen that yields lighter hydrocarbons on maturation than typical type ii kerogen. maturity gradients were constructed by converting the reflectance values of vitrinite-like particles to vitrinite-equivalent values following petersen et al. (2013) since vitrinite-like particles in the alum shale mature at lower temperatures than true vitrinite. the modelling showed that the lower palaeozoic shales were buried within a caledonian foreland basin and that large volumes of oil were probably generated during the silurian (gautier et al. 2013). in most areas, kerogen subsequently attained a maturation rank of dry gas, cracking the previously formed oil. in the carboniferous and early permian, the palaeozoic succession was faulted, tilted and subjected to intensive erosion (fig. 2). local permo-carbonifereous igneous intrusive rocks occur in the terne-1 borehole and elsewhere. however, these did not affect the regional maturity related to burial. in the area of the terne-1 borehole, subsidence resumed in the permo-triassic and maximum reburial probably occurred in cretaceous to early palaeogene time as is the general scenario in denmark (fig. 2). modelling suggests that the thermal rank reached during the palaeozoic was not exceeded during the reburial of the terne-1 area. nevertheless, because of sparse data and modelling uncertainty, we cannot exclude that some shale could have retained hydrocarbongeneration potential throughout the palaeozoic, and additional hydrocarbons may have formed during the mesozoic and cenozoic in some areas. the geological model for the assessment in denmark only two boreholes outside the skåne–bornholm area penetrate the alum shale (slagelse-1 and terne-1; fig. 1), hence the prospective area of the alum shale was delimited largely without borehole data. the analysis was based on maps of (1) the depth to the base of the palaeozoic (lassen & thybo 2012), (2) the distribution of palaeozoic strata (vejbæk & britze 1994) and (3) the regional thickness of the alum shale formation and its subdivisions (updated and somewhat modified from buchardt et al. 1997). these maps were used to identify areas where the alum shale is thicker than 20 m, gas mature and within a current depth interval of 1.5–7 km, which are relevant parameters for gas exploration. the prospective areas (fig. 3a) largely follow the margins of the norwegian–danish basin. alum shale is most likely also present in the central part of the basin, but the shale is here buried too deeply for exploration. the alum shale thins out or is missing on the ringkøbing–fyn high and southwards towards the caledonian front (fig. 1). sweet-spot mapping within the prospective area exploration undertaken by shell in skåne, southern sweden (boreholes a3, b2 and c3 in fig. 1), indicates that the alum shale formation, which is now located at 700–800 m depth, does not contain gas in economically producible quantities and that gas leakage from the shale has increased the risk for a viable gas play (pool et al. 2012). reservoir pressure reduction caused by uplift and loss of reservoir integrity due to faulting and fracturing are the likely mechanisms of gas loss. in skåne the gas may have leaked out through millions of years of uplift and progressive erosion since it formed more than 400 million years ago. in denmark, in contrast to skåne, the palaeozoic shale was reburied in the mesozoic and thus may retain gas to a greater degree. two types of areas with different risks of gas leakage were defined in the geological model, based on the thickness of palaeozoic strata mapped by lassen & thybo (2012). preserved thickness is taken as the best indicator for the magnitude of uplift and thus for the risk of reservoir depressurisafig. 2. timing of main events affecting the gas potential in the alum shale in southern scandinavia. palaeogene 2.6 quarternary neogene periodma era eon cretaceous jurassic triassic permian carboniferous devonian silurian ordovician cambrian precambrian pa lae oz oi c ph an er oz oi cm es oz oi c c en oz oi c 23 66 145 201 252 299 359 419 444 485 541 uplift uplift and possible loss of gas volcanic intrusions renewed burial burial and generation of gas deposition of the alum shale 21 tion in late palaeozoic time (fig. 3b). accordingly, within the prospective area, ‘sweet spots’ were defined as fault blocks that contain alum shale overlain by more than 1 km of palaeozoic strata (e.g. below the blue line in fig. 4), indicating less intensive late palaeozoic uplift and erosion and, hence, greater probability of gas retention (fig. 4). where the alum shale is overlain by less than 1 km of lower palaeozoic strata, the formation is inferred to have been uplifted to less than 1 km during the late palaeozoic, and those areas are therefore classified as non-sweet spot areas in the assessment. all the boreholes drilled so far in the alum shale in denmark and skåne play have been in non-sweet spots as defined here, with the highest reported gas saturation of 20% (pool et al. 2012). hence the quality of sweet spots remains to be tested. the difference in uplift history, and thus potentially in the gas content, is accounted for in the assessment model of gautier et al. (2013) by adopting different estimated ultimate recovery (eur) and success ratios for boreholes drilled in sweet spots (average eur 13.1 × 106 nm3 gas) versus nonsweet spots (average eur 6.7 × 106 nm3). development strategies the usgs assessment methodology assumes unrestricted application of best practice current technology, which in the present case is expected to be horizontal drilling with multistage hydrofracturing. in denmark the ordovician–silurian shale overlying the alum shale may constitute a rather thick (c. 300 m) additional interval in which other development strategies may be relevant. this inference is based on fig. 3. a: prospective areas in denmark for gas in the alum shale. b: distribution of sweet spots versus non-sweet spots within the prospective area of denmark (the term sweet spot is defined in the text). alum shale is likely also present in the deeper parts of the norwegian–danish basin, but here it is buried more than 7 km, i.e. too deeply for shale gas exploration with the current costs of drilling. alum shale non-sweet spot alum shale sweet spot alum shale buried 5.0–7.0 km alum shale buried 1.5–5.0 km 50 km a b norwegian–danish basin norwegian–danish basin 2222 the terne-1 borehole where a 250 m thick shale interval with toc values of 1–3% overlies the alum shale. these stratigraphic intervals are the targets for exploration in poland, lithuania and other countries in the eastern sector of the basin and may constitute an important additional reservoir in denmark. in addition, a tight gas play in upper silurian or lower permian sections may also be present in the subsurface of denmark and might add to the unconventional resource estimate. conclusions the estimated technically recoverable shale gas resource is comparable to the total volume of gas produced from the danish part of the north sea in the period 1972–2011 and twice the amount of remaining reserves of conventional gas in the danish sector of the north sea. however, in contrast to the resource estimates for the north sea, the estimated shale gas resource does not take economic viability into account. shale gas exploration in denmark is in its early stages. this is reflected in the large range of the estimate. it is thus crucial to obtain information from new boreholes, notably from sweet-spot areas, in order to calibrate and constrain the resource estimation model. the impact on the resource estimate from other development strategies or from additional play intervals and plays is not taken into consideration in the gas resource estimate by gautier et al. (2013). whether this is relevant awaits the evaluation of the first danish exploration borehole to be drilled in the lower palaeozoic in northern jylland. references buchardt, b., nielsen, a.t. & schovsbo, n.h. 1997: alun skiferen i skandinavien. geologisk tidsskrift 1997(3), 1–30. gautier, d.l., charpentier r.r., gaswirth, s.b., klett, t.r., pitman, j.k., schenk, c.j., tennyson, m.e. & whidden, k.j. 2013: undiscovered gas resources in the alum shale, denmark. u.s. geological survey fact sheet 2013–3103, 4 pp. lassen, a. & thybo, h. 2012: neoproterozoic and palaeozoic evolution of sw scandinavia based on integrated seismic interpretation. precambrian research 204–205, 75–104. mogensen, t.e. & korstgård, j.a. 2003: triassic and jurassic transtension along part of the sorgenfrei–tornquist zone, in the danish kattegat. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 439–458. nielsen, a.t. & schovsbo, n.h. 2011: the lower cambrian of scandinavia: depositional environment, sequence stratigraphy and palaeogeography. earth science reviews 107, 207–310. petersen, h.i., schovsbo, n.h. & nielsen, a.t. 2013: reflectance measurements of zooclasts and solid bitumen in lower palaeozoic shales, southern scandinavia: correlation to vitrinite reflectance. international journal of coal petrology 114, 1–18. pool, w., geluk, m., abels, j. & tiley, g. 2012: assessment of an unusual european shale gas play: the cambro-ordovician alum shale, southern sweden. proceedings of the society of petroleum engineers/european association of geoscientists and engineers unconventional resources conference, 20–22 march, 2012, vienna, 152339. schovsbo, n.h., nielsen, a.t., klitten, k., mathiesen, a. & rasmussen, p. 2011: shale gas investigations in denmark: lower palaeozoic shales on bornholm. geological survey of denmark and greenland bulletin 23, 9–12. vejbæk, o.v. & britze, p. (compilers) 1994: geological map of denmark 1:750 000. top pre-zechstein (two-way traveltime and depth). danmarks geologiske undersøgelse kortserie 45, 8 pp., 3 maps. authors’ addresses n.h.s, geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: nsc@geus.dk a.t.n., natural history museum of denmark, øster voldgade 5–7, dk-1350 copenhagen k, denmark. d.l.g., consulting geologist, 3954 nelson court, palo alto, california 94306, usa. fig. 4. conceptual cross section showing the subsurface geology in northern jylland. the alum shale is present in tilted fault blocks below the caledonian unconformity. sweet spots are the areas with lowest risk of gas leakage during late palaeozoic uplift and are defined as regions where the alum shale is overlain by more than 1 km of lower palaeozoic strata. modified and generalised from mogensen & korstgård (2003). sæby-1 0 4 2 6 ns d ep th (k m ) alum shale silurian shale ordovician– clay and sandstone unconformity 1 km depth in palaeozoic fault sweet spot basement 5 km palaeozoic post-palaeozoic clay and siltstone chalk sandstone geological survey of denmark and greeland bulletin 4, 2003, pp 61-64 61 following the expected ratification in 2004 of the united nations convention on the law of the sea (unclos from 1982), denmark, greenland and the faroe islands have a period of maximum 10 years to make claims beyond 200 nautical miles (nm) in five potential areas off greenland and the faroe islands (fig. 1). in order to provide the necessary database, the danish continental shelf project has been launched by the ministry for science, technology and innovation in cooperation with the faroese and greenland home rule governments. several institutions are participating in this project, with the geological survey of denmark and greenland (geus) as the coordinator of the technical work for the greenland part of the project, and sharing the responsibility for coordination of the faroese part with the faroese geological survey (jfs). background article 76 of unclos is the key to future jurisdiction over resources on and below the seabed beyond 200 nm. the right to explore and exploit these resources, which include both non-living resources (hydrocarbons and minerals) and bottom-dwelling living resources, may have significant economic implications. furthermore, jurisdiction of the extended continental shelf gives the right to regulate transport, environment and research. the technical data needed for a submission to the commission on the limits of the continental shelf (clcs) include geodetic, bathymetric, geophysical and geological data. the most critical issue is to be able to demonstrate a natural prolongation of the submerged land territory based on geological and geomorphological factors, and later to document claims in detail by using the various formulas and constraint lines of article 76 (commission on the limits of the continental shelf (clcs) 1999; cook & carleton 2000). for construction of these lines it is necessary to know distances from territorial sea base lines (+ 200 and 350 nm), to define the location of the foot of the continental slope and the 2500 m isobath, and to know the sediment thickness beyond the foot of the slope. the foot of the continental slope (fos) is defined as the point of maximum change of gradient at the base of the continental slope. areas of interest around greenland there are three potential claim areas off greenland. one south of greenland is outlined by the 200 nm limit, a yet to be established boundary with canada, and the new outer limit. the second area north-east of greenland is outlined by the 200 nm line from greenland, jan mayen (norway) and svalbard (norway) and the new outer limit. the third is north of greenland, outlined by the 200 nm limit and yet to be established boundaries with canada and possibly also russia and/or norway, and a new outer limit. all three areas are situated along margins of mixed rifted and strike-slip nature, and contain significant successions of volcanic and sedimentary rocks. geological survey of denmark and greenland bulletin 4, 61–64 (2004) © geus, 2004 exploring for extended continental shelf claims off greenland and the faroe islands – geological perspectives christian marcussen, flemming g. christiansen, trine dahl-jensen, martin heinesen, steen lomholt, jens jørgen møller and kai sørensen fig. 1. map of the north atlantic region. arrows indicate the five potential claim areas of interest. 1: lomonosov ridge, 2: amundsen basin, 3: morris jesup rise, 4: gakkel ridge, 5: knipovitch ridge, 6: east greenland ridge, 7: mohns ridge, 8: eiriks ridge, 9: labrador sea, 10: hatton bank, 11: rockall bank, 12: rockall trough. jm, jan mayen. south of greenland the eiriks ridge is assumed to be a natural prolongation of southern greenland with the foot of slope on the deep-water side of the ridge (fig. 2). the existence of very thick sedimentary successions between greenland and canada, especially within the extinct spreading zone, may form the basis for a claim far out into the labrador sea. there is a general consensus on a tectonic model with sea-floor spreading in the labrador sea in paleocene–eocene time, possibly continuing into the miocene. models have changed considerably with time, since the early work by srivastava and co-workers suggested large areas were underlain by oceanic crust of late cretaceous to miocene age, anomaly 33–20 time (e.g. srivastava 1978; roest & srivastava 1989). these models were revised by chalmers & pulvertaft (2001), who suggest spreading from anomaly 27–20 time. the 2003 geophysical programme focused on data acquisition along the eiriks ridge, and within and across the extinct spreading zone, to document and correlate thick sedimentary successions. a total of 1500 km reflection seismic data were acquired. the new seismic data confirm the sedimentary nature of the eiriks ridge (fig. 2) and furthermore show that sediment thicknesses are a potential factor for a future claim beyond 200 nm. north-east of greenland the east greenland ridge is assumed to be a natural prolongation of north-eastern greenland, and the foot of the slope extends around the ridge (fig. 3). the thick sedimentary suc62 fig. 3. a: map of the east greenland ridge region. the seismic lines acquired in 2002 are shown in orange and white, the 200 nm limits of greenland, svalbard and jan mayen in red, and the 350 nm limit of greenland in blue. the stippled black line is the unofficial median line between greenland and svalbard; the full black line is the official border between greenland and norway (jan mayen), and greenland and iceland. b: seismic section across the east greenland ridge. position indicated in fig. 3a (white line). fig. 2. a: map of the south greenland region showing seismic lines acquired in 2003 (orange and white) and older seismic lines (black). stippled line is the unofficial median line with canada; 200 nm line in red. b: seismic line along the eiriks ridge. position indicated in fig. 2a (white line). cessions that exist both north and south of the ridge are likely to contribute to the claim area. although there is general consensus on the opening history of the north atlantic, problems with details of the structural elements are apparent, especially north of the east greenland ridge where major strike-slip movements occur in the region between the shelf break and the knipovitch ridge. a simpler spreading pattern is observed south of the ridge since anomaly 24 b time, with active spreading along the mohns ridge (mosar et al. 2002; tsikalas et al. 2002). in the summer of 2002 geus and the university of bergen carried out a joint refraction and reflection seismic survey over the east greenland ridge. there is a marked difference in water depth and subsurface structure north and south of the ridge (fig. 3). the preliminary wide-angle velocity model shows standard oceanic crust to the south, while the northern side of the ridge and the crust further north may be stretched continental crust. north of greenland the lomonosov ridge and the morris jesup rise are assumed natural prolongations of northern greenland. if relatively thick sedimentary successions can be demonstrated in the amundsen basin it may be possible to enlarge the potential claim area (fig. 1). the existing data coverage from the arctic ocean north of greenland is very sparse, due to the adverse physical conditions with metre-thick sea ice and many pressure ridges. there is, however, general consensus on a tectonic model with active spreading since paleocene time along the gakkel ridge, and with the lomonosov ridge most likely consisting of continental crust separated from the barents–kara shelf (jokat et al. 1995; lawver et al. 2002). preliminary studies in 2004 will focus on testing data acquisition methods on the sea ice, and on a passive earthquake seismological experiment for the crustal structure of north greenland. subsequent studies will include acquisition of refraction seismic data on the sea ice along the innermost parts of the lomonosov ridge, followed by data acquisition from ice breakers across the lomonosov ridge and in the amundsen basin. areas of interest around the faroe islands the faroe islands consist of basaltic rocks with a cumulative stratigraphic thickness of more than 6.5 km resting on top of presumed precambrian basement (ellis et al. 2002). during the initial phases of continental break-up between europe and greenland, the faroe islands and the hatton–rockall area (the faroe–rockall plateau) were partly isolated from the main european continent. subsequent shift of the break-up axis to the west of the plateau resulted in extensive (basaltic) volcanism, seafloor spreading and the creation of the northeast atlantic ocean between europe and greenland. the two potential claim areas off the faroe islands are an area northeast of the islands, and the hatton–rockall area to the southwest (fig. 4). 63 fig. 4. a: map of the faroe islands region. the designated area is outlined with blue lines. the red lines indicate the 200 nm limits of the surrounding coastal states. continental shelf median lines are drawn in black. the position of the three seismic lines acquired in 2003 are shown in orange and white. b: seismic line los_fo_03-1. position indicated in fig. 4a (white line). the north-eastern area the basaltic rocks that form the faroe islands reach far offshore on the continental shelf and continue onto the continental margins to the north of the islands (boldreel & andersen 1994). in the central part of this region the slope beyond the shelf break is steep and the transition from shelf to deep ocean sea floor is narrow. elsewhere, the slope is more gentle (fig. 4). thick sediment accumulations in the deep-water areas favour claims extending beyond the 200 nm limit. three seismic lines were acquired in 2003 (fig. 4) to study the continent–ocean transition (cot) and to assess the sediment thickness along the continental shelf margin. the south-western area based on unclos, article 76, and on the assumption that the faroe–rockall plateau constitutes a micro-continent, the danish authorities designated a large continental shelf area to the south-west of the faroe islands in 1985 (fig. 4). this area includes the parts of the hatton and rockall banks situated outside the 200 nm limits of the neighbouring states (great britain, ireland and iceland), who in their turn have made individual designations for the same area. the basaltic rocks that form the faroe islands continue and thin south-westwards, and disappear at several locations at the hatton and rockall banks where the underlying rocks are exposed at seabed. the plateau margin to the west is relatively simple with a well-defined slope area (fig. 4). different volcanic and tectonic features to the south of the plateau complicate the marginal area in this region. towards the east the plateau borders the rockall trough, which reaches water depths of 3–4 km in its southernmost part. work planned for 2004 includes a deep reflection and refraction seismic programme and a geochemical study. conclusion the danish continental shelf project has so far acquired new data in three out of five potential claim areas off greenland and the faroe islands, and much more data acquisition and interpretation will follow in the next few years. the results of the project, together with similar projects by neighbouring countries, will create new focus on the geology and tectonics of the north atlantic and arctic regions. references boldreel, l.o. & andersen, m.s. 1994: tertiary development of the faroe–rockall plateau based on reflection seismic data. bulletin of the geological society of denmark 41(2), 162–180. chalmers, j.a. & pulvertaft, t.c.r. 2001: development of the continental margins of the labrador sea – a review. in: wilson, r.c.l. et al. (eds): non-volcanic rifting of continental margins: a comparison of evidence from land and sea. geological society special publication (london) 187, 77–105. commission on the limits of the continental shelf (clcs) 1999: scientific and technical guidelines of the commission on the limits of the continental shelf. clcs/11, 91 pp. + annexes. cook, p.j. & carleton, c.m. (eds) 2000: continental shelf limits: the scientific and legal interface, 363 pp. new york: oxford university press. ellis, d., bell, b.r., jolley, d.w. & o’callaghan, m. 2002: the stratigraphy, environment of eruption and age of the faroes lava group, ne atlantic ocean. in: jolley, d.w. & bell, b.r. (eds): the north atlantic igneous province: stratigraphy, tectonic, volcanic and magmatic processes. geological society special publication (london) 197, 253–269. jokat, w., weigelt, e., kristoffersen, y., rasmussen, t. & schöne, t. 1995: new insights into the evolution of the lomonosov ridge and the eurasian basin. geophysical journal international 122(2), 378–392. lawver, l.a., grantz, a. & gahagan, l.m. 2002: plate kinematic evolution of the present arctic region since the ordovician. in: miller, e.l., grantz, a. & klemperer, s.l. (eds): tectonic evolution of the bering shelf–chukchi sea–arctic margin and adjacent landmasses. geological society of america, special paper 360, 333–358. mosar, j., lewis, g. & torsvik, t.h. 2002: north atlantic sea-floor spreading rates; implications for the tertiary development of inversion structures of the norwegian–greenland sea. journal of the geological society (london) 159(5), 503–515. roest, w.r. & srivastava, s.p. 1989: sea-floor spreading in the labrador sea: a new construction. geology 17, 1000–1003. srivastava, s.p. 1978: evolution of the labrador sea and its bearing on the early evolution of the north atlantic. geophysical journal of the royal astronomical society 52, 313–357. tsikalas, f., eldholm, o. & faleide, j.i. 2002: early eocene sea floor spreading and continent–ocean boundary between jan mayen and senja fracture zones in the norwegian–greenland sea. marine geophysical researches 23(3), 247–270. authors’ addresses c.m., f.g.c., t.d.-j., s.l., j.j.m. & k.s., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: cma@geus.dk m.h., jar›frø›isavni› (jfs) – faroese geological survey, postboks 3169, fo-110 tórshavn, the faroe islands. 64 << /ascii85encodepages false /allowtransparency false /autopositionepsfiles true /autorotatepages /all /binding /left /calgrayprofile (dot gain 20%) /calrgbprofile (srgb iec61966-2.1) /calcmykprofile (u.s. web coated \050swop\051 v2) /srgbprofile (srgb iec61966-2.1) /cannotembedfontpolicy /warning /compatibilitylevel 1.4 /compressobjects /tags /compresspages true /convertimagestoindexed true /passthroughjpegimages true /createjdffile false /createjobticket false /defaultrenderingintent /default /detectblends true /colorconversionstrategy /leavecolorunchanged /dothumbnails false /embedallfonts true /embedjoboptions true /dscreportinglevel 0 /syntheticboldness 1.00 /emitdscwarnings false /endpage -1 /imagememory 1048576 /lockdistillerparams false /maxsubsetpct 100 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<feffd5a5c0c1b41c0020c778c1c40020d488c9c8c7440020c5bbae300020c704d5740020ace0d574c0c1b3c4c7580020c774bbf8c9c0b97c0020c0acc6a9d558c5ec00200050004400460020bb38c11cb97c0020b9ccb4e4b824ba740020c7740020c124c815c7440020c0acc6a9d558c2edc2dcc624002e0020c7740020c124c815c7440020c0acc6a9d558c5ec0020b9ccb4e000200050004400460020bb38c11cb2940020004100630072006f0062006100740020bc0f002000520065006100640065007200200035002e00300020c774c0c1c5d0c11c0020c5f40020c2180020c788c2b5b2c8b2e4002e> /chs <feff4f7f75288fd94e9b8bbe7f6e521b5efa76840020005000440046002065876863ff0c5c065305542b66f49ad8768456fe50cf52068fa87387ff0c4ee563d09ad8625353708d2891cf30028be5002000500044004600206587686353ef4ee54f7f752800200020004100630072006f00620061007400204e0e002000520065006100640065007200200035002e00300020548c66f49ad87248672c62535f003002> /cht <feff4f7f752890194e9b8a2d5b9a5efa7acb76840020005000440046002065874ef65305542b8f039ad876845f7150cf89e367905ea6ff0c4fbf65bc63d066075217537054c18cea3002005000440046002065874ef653ef4ee54f7f75280020004100630072006f0062006100740020548c002000520065006100640065007200200035002e0030002053ca66f465b07248672c4f86958b555f3002> >> >> setdistillerparams << /hwresolution [2400 2400] /pagesize [595.000 842.000] >> setpagedevice geological survey of denmark and greenland bulletin 1, 813-863 813 shallow marine syn-rift sedimentation: middle jurassic pelion formation, jameson land, east greenland michael engkilde and finn surlyk the middle jurassic pelion formation – fossilbjerget formation couplet of jameson land, east greenland, is a well-exposed example of the middle jurassic inshore–offshore successions characteristic of the rifted seaways in the northwest european – north atlantic region. early jurassic deposition took place under relatively quiet tectonic conditions following late permian – earliest triassic and early triassic rift phases and the lower jurassic stratal package shows an overall layer-cake geometry. a long-term extensional phase was initiated in middle jurassic (late bajocian) time, culminated in the late jurassic (kimmeridgian–volgian), and petered out in the earliest cretaceous (valanginian). the upper bajocian – middle callovian early-rift succession comprises shallow marine sandstones of the pelion formation and correlative offshore siltstones of the fossilbjerget formation. deposition was initiated by southwards progradation of shallow marine sands of the pelion formation in the late bajocian followed by major backstepping in bathonian–callovian times and drowning of the sandy depositional system in the middle–late callovian. six facies associations are recognised in the pelion–fossilbjerget couplet, representing estuarine, shoreface, offshore transition zone and offshore environments. the north–southtrending axis of the jameson land basin had a low inclination, and deposition was sensitive to even small changes in relative sea level which caused the shorelines to advance or retreat over tens to several hundreds of kilometres. eight composite sequences, termed p1–p8, are recognised and are subdivided into a total of 28 depositional sequences. the duration of the two orders of sequences was about 1–2 ma and 360,000 years, respectively. the upper bajocian p1–2 sequences include the most basinally positioned shallow marine sandstones, deposited during major sealevel lowstands. the lowstands were terminated by significant marine flooding events, during which sandstone deposition was restricted to northern, more proximal parts of the basin. the upper bajocian – middle bathonian p3–4 sequences show an overall progradational stacking pattern. the sequence boundary at the top of p4 marks a significant shift in stacking pattern, and the upper bathonian – middle callovian p5–8 sequences show large-scale backstepping, terminating in a widespread condensed succession at the distal, southern end of the basin. the largescale backstepping was governed by combined tectonically-induced subsidence, reflecting increased rates of extension, and eustatic sea-level rise. the depositional trends of the pelion formation – fossilbjerget formation couplet provide a well-exposed analogue to contemporaneous subsurface deposits which form major hydrocarbon reservoirs on the west norway shelf, and in the northern north sea. keywords: east greenland, jameson land, upper bajocian – middle callovian, pelion formation, fossilbjerget formation, sedimentology, sequence stratigraphy, shallow marine – offshore environments, regressive–transgressive clastic wedge m.e.* & f.s., geological institute, university of copenhagen, geocenter copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: finns@geo.geol.ku.dk *present address: denerco oil a/s, kongevejen 100c, postbox 110, dk-2840 holte, denmark. e-mail: mke@denerco.dk geological survey of denmark and greenland bulletin 1, 813–863 (2003) © geus, 2003 the middle jurassic of east greenland was characterised by the onset of rifting in the late bajocian, major reorganisation of drainage systems, and a high influx of sand. a gradual westand northwards onlap onto progressively older strata took place during late bajocian – callovian times. upper bajocian sandand siltstones of the basal pelion formation rest on lower bajocian mudstones of the sortehat formation in jameson land at the southern end of the east greenland basin (fig. 1); lithostratigraphic usage in this paper follows the provisional revised scheme in surlyk (2003, this volume, fig. 5). the age of the sortehat formation is determined by dinoflagellates (underhill & partington 1994; koppelhus & hansen 2003, this volume), and by sr-isotope stratigraphy (m. engkilde, unpublished data). further north, on traill ø, presumed middle jurassic fluvial pebbly sandstones of the bristol elv formation (therkelsen & surlyk in press) or upper bajocian shallow marine sandstones of the pelion formation rest on upper triassic redbeds of the flemming fjord formation and there is no evidence for the presence of lower jurassic strata. at hold with hope, the pelion formation overlies the lower triassic wordie creek formation (stemmerik et al. 1997; vosgerau et al. in press a), and on wollaston forland it rests on a thin upper permian carbonate–evaporite unit or directly on caledonian crystalline basement. further north again, on kuhn ø, upper bathonian or callovian strata of the fluvial bastians dal formation (alsgaard et al. 2003, this volume) or the marine pelion formation rest directly on the basement. in hochstetter forland, callovian coal-bearing paralic deposits of the upper bathonian(?) – middle callovian muslingebjerg formation or marine callovian–oxfordian sandstones of the payer dal formation onlap the caledonian basement. the progressive northwards truncation of the underlying succession and younging of the base of the middle jurassic strata probably reflect late early jurassic development of a rift dome in north-east greenland and subsequent middle jurassic domal erosion, deflation and transgression (surlyk 1977a, 1978) similar to the roughly contemporaneous dome in the north sea (sellwood & hallam 1974; whiteman et al. 1975; eynon 1981; ziegler 1988; underhill & partington 1993, 1994). the reality of early jurassic uplift has been confirmed on the basis of fission track thermochronology by johnson & callagher (2000). in contrast, the jameson land area shows no evidence for early jurassic uplift and cooling (mathiesen et al. 2000). the mesozoic succession was uplifted during the neogene and is well-exposed, notably in jameson land, milne land, traill ø, wollaston forland, and kuhn ø (figs 1, 2). in this study, the focus is on the upper bajocian – middle callovian pelion and fossilbjerget formations of jameson land which form a northand westwards thickening wedge of shallow marine sand814 milne land staunings alper traill ø geographical society ø clavering ø kuhn ø store koldewey 100 km 22°w 26°w 24°w 20°w 18°w 16°w 26°w 22°w24°w28°w 72°n 74°n 76°n fault (indeterminate type) inferred fault normal fault reverse fault hochstetter forland wollaston forland hold with hope kejser franz joseph fjord scoresby sund kong oscar fjord jameson land liverpool land greenland fig. 1. map showing fault zones active during the mesozoic in east greenland. place names used in the text are shown. modified from surlyk (1991). 815 stones and siltstones (figs 1–3; surlyk et al. 1973; heinberg & birkelund 1984; engkilde & surlyk 1993; engkilde 1994; surlyk & noe-nygaard 2000; larsen et al. 2003, this volume; surlyk 2003, this volume). the region has a long history of investigation and previous work on the middle jurassic succession of jameson land was largely focused on the general stratigraphy (madsen 1904; koch 1929, 1950; rosenkrantz 1929, 1934, 1942; spath 1932, 1947; stauber 1940; callomon 1959, 1961, 1970, 1972, 1993, 1994; birkelund et al. 1971; surlyk & birkelund 1972; surlyk et al. 1973, 1993; birkelund & perch-nielsen 1976; surlyk 1991). the only detailed study of the middle jurassic succession was by heinberg & birkelund (1984), who analysed the trace fossil assemblages and facies patterns of the pelion and fossilbjerget formations in jameson land. the low gradient physiography, and general shallow water depth of the basin during middle jurassic times, made deposition very sensitive to even small changes in relative sea level. such changes caused the shorelines to advance or retreat over tens to several hundreds of kilometres, with the formation of laterally extensive depositional units. the interpretation of changes in spatial distribution of the units through time is controlled by a detailed ammonite biostratigraphy and large-scale depositional variations and geometries. the aim of the study is to establish depositional models and a sequence stratigraphic framework for the middle jurassic succession of jameson land. the architecture and facies trends through time of the pelion– fossilbjerget couplet provide a well-exposed and welldated depositional analogue to contemporaneous major hydrocarbon reservoirs on the west norway shelf and the northern north sea. upper jurassic – lower cretaceous middle jurassic upper triassic – lower jurassic devonian – middle triassic crystalline rocks normal fault 1 goniomyakløft 2 fortet 3 ugleelv 4 rævekløft 5 mikael bjerg 6, 6a trefjord bjerg (s, sw) 7 trefjord bjerg (n) 8 olympen (s) 9 olympen (e) 10 pelion/parnas 11 pelion (n) 12 pothorst bjerge 13 claudius clavus bjerge 14 antarctic havn 70°30′n 50 km 25°w 24°w 23°w 22°w 71°30′n 71°n 5 10 6 6a 7 98 1 2 3 4 12 11 14 13 a′ a 0 km 10 160 km140120100806040200 20 a a′ crystalline basement carboniferous – lower permian upper permian – jurassicdevonian jameson land liverpool land nw se m fig. 2. geological sketch map of jameson land. numbers refer to localities described in the text. the section a–a′′ is shown in fig. 3. modified from surlyk et al. (1973). fig. 3. nw–se section of central jameson land (for location, see fig. 2). the section shows the faulted, eastwards rising crystalline basement of liverpool land. the mesozoic sediments onlapped the liverpool land crystalline basement or permian sediments towards the east before removal by modern erosion. modified from larsen & marcussen (1992); m, moho. 816 geological setting the late palaeozoic – mesozoic extensional basins of east greenland are exposed over c. 800 km in a south– north direction, from jameson land and milne land in central east greenland to store koldewey in north-east greenland (fig. 1). the mesozoic deposits were largely deposited in two major basins, the jameson land and the wollaston forland basins. the jameson land basin is c. 140 km wide in south jameson land and milne land, narrowing northwards from traill ø to geographical society ø and hold with hope (figs 1, 3). the wollaston forland basin is c. 50 km wide in the clavering ø – wollaston forland area, and narrows northwards through kuhn ø and hochstetter forland (fig. 1). marine jurassic sandstones are found at store koldewey, and represent the western edge of a mesozoic basin which continues offshore (fig. 4; surlyk et al. 1981, 1986; surlyk 1990a). the east greenland basins formed the westernmost segment of the major north–south-oriented system of rift basins situated between greenland and norway (fig. 5; ziegler 1988; larsen 1990; doré 1991, 1992; surlyk et al. 1993). at least 5 km of upper permian – mesozoic sediments were deposited in the jameson land basin (fig. 3). the late permian to cretaceous basin evolution was characterised by post-rift thermal conocean-to-continent transition inferred faults precambrian and caledonian crystalline basement upper palaeozoic sediments mesozoic sediments inferred offshore rift basins shallow depth or outcropping lower tertiary basalts jameson land 200 km 10°w 0°30°w 20°w 75°n 70°n 20°w 75°n 80°n 70°n 30°w 40°w fig. 4. map showing the position of inferred mesozoic sedimentary basins below the north-east greenland shelf, based on aeromagnetic surveys and seismic data. modified from larsen (1990). land margin areas major fault, down-throw indicated 60°n 55°n 50°n pa lae ol at itu de 500 km wollaston forland basin jameson land basin in tr arif t h igh greenland laurentian shield sweden norway baltic shield fig. 5. palaeogeographic map of the north sea and northern north atlantic region during the middle jurassic, showing the position of the basins in east greenland, the norwegian shelf and in the north sea. modified from doré (1991). traction and sediment loading, following late palaeozoic rifting, interrupted by rift events in late permian – early triassic, early triassic, middle–late jurassic, early and late cretaceous times (donovan 1953; callomon 1972; surlyk et al. 1973; surlyk 1977b, 1990a, 1991, 2003, this volume; clemmensen 1980; surlyk & clemmensen 1983; larsen 1990; larsen & marcussen 1992; price & whitham 1997; surlyk & noe-nygaard 2001a). major regional uplift of the areas north of jameson land took place in early jurassic times, probably close to the early–middle jurassic transition. it was associated with changes in basin configuration and development of new drainage and transport patterns. the uplift has been related to the formation of a large rift dome north of jameson land, analogous to the north sea dome developed at the triple junction between the central graben, the viking graben and the witch ground graben (surlyk 1977a, 1978). the uplift was marked by the sudden influx mainly from the north of large volumes of coarse-grained sediments in the northwards thickening succession (surlyk et al. 1973, 1981; surlyk & clemmensen 1983; heinberg & birkelund 1984; engkilde & surlyk 1993; engkilde 1994). similar middle jurassic uplift, succeeded by deposition of large volumes of coarse clastic deposits, is also documented from the mid-norway shelf and the north sea (doré 1991, 1992). these deposits now form many of the most productive oil reservoirs in the north sea and at the mid-norway margin (doré 1991; mitchener et al. 1992; cordey 1993). the eastern margin of the jameson land basin, represented by the present-day liverpool land, was a structural high in mesozoic times that was periodically exposed and accumulated only a thin sediment cover (figs 2, 3; rosenkrantz 1942; birkelund & perch-nielsen 1976; birkenmajer 1976; clemmensen 1980; surlyk et al. 1981; dam & surlyk 1993, 1998). by middle jurassic times, north-westerly and westerly sediment sources were dominant (bromley et al. 1970; surlyk et al. 1973; birkelund & perch-nielsen 1976; surlyk 1977b, 1991; callomon & birkelund 1980). fluvial and deltaic sediments occur at the base of the succession on traill ø and geographical society ø, and further north on kuhn ø and hochstetter forland (price & whitham 1997; stemmerik et al. 1997; alsgaard et al. 2003, this volume; therkelsen & surlyk in press). the nature of the southern basin margin is not known, but seismic data suggest that jurassic deposits extend south of scoresby sund (larsen & marcussen 1992). the middle jurassic basin fill onlaps precambrian and caledonian crystalline basement along the western basin margin in milne land (callomon & birkelund 1980; larsen & marcussen 1992; larsen et al. 2003, this volume). the northwards termination of the jurassic liverpool land basement high is not precisely known. an eastwards tilted fault-block was formed by middle jurassic rifting in south-eastern traill ø, and eastwards dips are also noted in north-east jameson land. this is opposite to all other jurassic fault-blocks in east greenland which show marked westwards dips and suggest direct marine communication towards the east in this area (carr 1998; vosgerau et al. in press b). an important tectonic zone is situated in northern jameson land, separating the broad jameson land platform from the northern region which is characterised by late jurassic and cretaceous tilted fault blocks. the cross fault zone probably coincides with a zone of nw–se-trending deep-seated faults recognised by dam et al. (1995). the fault zone was reactivated during cenozoic basin uplift (vischer 1943; donovan 1953; haller 1971; surlyk 1977b; surlyk et al. 1981, 1993; larsen & marcussen 1992). stratigraphy of the pelion and fossilbjerget formations the pelion and fossilbjerget formations (lower vardekløft group) form a north-westwards thickening wedge of late bajocian – middle callovian age, covering about 10 ma (fig. 6). the formations overlie the aalenian – lower bajocian sortehat formation of the neill klinter group in jameson land (surlyk et al. 1973; dam & surlyk 1998; surlyk 2003, this volume, fig. 5). the base of the pelion–fossilbjerget couplet is a major sequence boundary, marked by a dramatic increase in the influx of coarse clastic sediments (surlyk 1991, 2003, this volume; surlyk et al. 1993; dam & surlyk 1998). farther north, on traill ø, the pelion formation overlies upper triassic redbeds of the flemming fjord formation (clemmensen 1980; fig. 6). in milne land, at the western basin margin, middle jurassic sandstones of the charcot bugt formation onlap crystalline basement (callomon & birkelund 1980; larsen et al. 2003, this volume). the top of the pelion–fossilbjerget couplet is a widespread condensed unit in southern jameson land. it is overlain by the upper callovian – middle oxfordian marine sandstones and mudstones of the olympen formation in northern jameson land and traill ø. in southern jameson land, the formations are overlain by condensed, distal deep-water mudstones of the olympen formation, followed by black mudstones and massive sandstones of the upper oxfordian – lower volgian hareelv formation (surlyk 1987, 2003, this volume; 817 surlyk & noe-nygaard 1998, 2001b; larsen & surlyk 2003, this volume). the pelion–fossilbjerget couplet thickens along the basin axis from 150 m in southern jameson land to 475 m in central and northern jameson land (surlyk 1977b; heinberg & birkelund 1984). the pelion formation thickens in the same direction from c. 50 m to about 400 m. the very high thickness value, close to 700 m, reported for the pelion formation at antarctic havn in northernmost jameson land (heinberg & birkelund 1984), probably includes the upper callovian – middle oxfordian olympen formation. if this very poorly exposed and strongly faulted area is excluded, figure 25 of heinberg & birkelund (1984) shows uniform thicknesses for all time slices from mikael bjerg in southern central jameson land and further north. the only significant decrease in thickness thus seems to take place in the region of sandstone pinch-out from mikael bjerg and further south. the pelion formation consists mainly of shallow marine sandstones, which are overlain by, and pass southwards into deeper-water, offshore siltstones of the fossilbjerget formation (fig. 6). the boundary between the two formations is strongly diachronous, and youngs 818 ns pe lio n fm a nd f os si lb je rg et f m c on de ns ed s ec tio n ju ra ss ic c al lo vi an b at ho ni an b aj oc ia n 173 169.2 164.4 161 ma chronostratigraphy boreal province chronozones subboreal and submediterranean province c. borealis c. indistinctus c. pompeckji a. arcticus a. greenlandicus a. ishmae a. cranocephaloide c. variabile c. calyx c. apertum n. subfurcatum g. garantiana p. parkinsoni z. zigzag p. progracilis p. hodsoni o. orbis c. discus m. herveyi p. koenigi s. calloviense k. jason e. coronatum c. nordenskjoeldi l u l m u l m u onlap p. athleta shallow marine sandstones offshore siltstones, mudstones hiatus pebbles a. tenuiplicatus t. subcontractus m. morrisi 200 km p8p8 p7p7 p6p6 p5p5 p4p4 p3p3 p2p2 p1p1 p2p2 p1 jameson land sortehat fm triassic fig. 6. chronostratigraphical scheme of the pelion and fossilbjerget formations in jameson land (time-scale from gradstein et al. 1994). p1–8 indicate composite depositional sequences. the number of high-order depositional sequences within the composite sequences is indicated by the saw-tooth pattern. the bulk sandstone parts of the composite sequences are late highstand deposits throughout central and northern jameson land. thick deposits of the transgressive and early highstand systems tracts are predicted to exist in western and northern basin margin areas; such deposits are not shown in this figure, but indicated schematically on fig. 36. two lowstand units are shown in sequences p1 and p2. it should be noted that precise correlation to the bajocian and bathonian stages of europe is not yet possible, due to faunal provincialism; the jurassic ammonite zonation is from callomon (1993). the figure is based on fig. 2 in surlyk (1991), with the addition of new data on the sequence stratigraphy. towards the north (callomon 1959, 1993; surlyk et al. 1973; heinberg & birkelund 1984). the formations show a record of cyclic regressions and transgressions on several orders, and the boundary between the formations is thus not one continuous surface, but can be considered a series of shingled marine flooding surfaces, which cap individual progradational shallow offshore – shoreface units of the pelion formation. the geometry of the formations, large-scale textural gradients, and the observed palaeocurrent patterns, indicate that the main sediment influx was from the north and northwest, and sediment transport was mainly towards the south, along the basin axis. the age relationships of the formations are based on a detailed ammonite biostratigraphy (spath 1932; callomon 1959, 1993; surlyk et al. 1973). the pelion– fossilbjerget couplet contains 16 ammonite zones, which are subdivided into 37 ammonite faunal horizons (figs 6, 7). correlation with the european ammonite zonation in the bajocian–bathonian is not possible due to faunal provincialism (see callomon 2003, this volume), and the precise age of the lower boundary of the pelion formation is uncertain but is tentatively placed in the earliest late bajocian (fig. 6, left part; callomon 1959, 1972, 1993; surlyk et al. 1973). this is corroborated by 87sr/86sr values which suggest an early late bajocian age for the base of the pelion formation, by comparison with the strontium isotope curve of jones et al. (1994) (m. engkilde, unpublished data). the sedimentary organic content is dominantly terrestrially derived and poorly preserved. abundant terrestrial pollen and spores occur throughout the formations together with subordinate marine dinoflagellate cysts that are abundant only at certain levels. spores and pollen have contributed little to the biostratigraphic subdivision of the formations as most middle and late jurassic species are long-ranging (m.d. muir in: sarjeant 1972; lund & pedersen 1985). dinoflagellate cysts have a higher stratigraphic potential which is not yet fully exploited (sarjeant 1972; smelror 1988; larsen et al. 2003, this volume). marine connections between the tethyan and boreal realms were re-established in the callovian, and good correlations of the ammonite successions exist for the callovian–kimmeridgian interval (fig. 6; callomon 1959, 1972, 1993; surlyk et al. 1973; birkelund & perch-nielsen 1976; birkelund et al. 1984; birkelund & callomon 1985). the pelion–fossilbjerget couplet consists of eight composite sequences which are subdivided into 28 simple coarsening-upwards sequences (following the definition of mitchum & van wagoner 1991). the duration of the composite sequences is one to a few million years, and several hundred thousand years for the simple sequences (fig. 6). the pelion and fossilbjerget formations form the lower part of the long-term regressive–transgressive vardekløft group (surlyk & noe819 a. glosense c. tenuiserratum c. densiplicatum c. cordatum q. mariae q. lamberti p. athleta e. coronatum k. jason s. calloviense p. koenigi c. nordenskjoeldi c. apertum c. calyx c. variabile a. cranocephaloide a. ishmae a. greenlandicus a. arcticus c. pompeckji c. indistinctus c. borealis decipia cf. decipiens cardioceras cf. densiplicatum cardioceras alphacordatum quenstedtoceras woodhamense longaeviceras keyserlingi kosmoceras cf. or aff. jason sigaloceras calloviense kepplerites galilaeii chamoussetia phillipsi kepplerites cf. gowerianus cadoceras septentrionale cadoceras nordenskjoeldi β cadoceras nordenskjoeldi α cadoceras cf. or aff. breve kepplerites tenuifasciculatus cadoceras apertum γ cadoceras apertum β cadoceras apertum α kepplerites vardekloeftensis kepplerites peramplus kepplerites rosenkrantzi kepplerites inflatus kepplerites tychonis arcticoceras cranocephaloide arcticoceras crassiplicatum arcticoceras ishmae β arcticoceras ishmae α arcticoceras harlandi arctocephalites freboldi arctocephalites greenlandicus arctocephalites micrumbilicatus arctocephalites delicatus arctocephalites arcticus cranocephalites episcopalis cranocephalites pompeckji cranocephalites furcatus cranocephalites carlsbergensis cranocephalites maculatus cranocephalites gracilis cranocephalites intermissus cranocephalites indistinctus α and β cranocephalites borealis α and β o xf or di an u m l u m l ba jo ci an ( pa rs ) ba th on ia n c al lo vi an chronozones faunal horizons fig. 7. middle jurassic faunal horizons and ammonite zones of east greenland. based on callomon (1993). nygaard 2000; surlyk 2003, this volume). in the jameson land basin, the age of the composite sequences is wellestablished to ammonite zone or subzone level. many of the simple sequences are also biostratigraphically dated by ammonites. sedimentology of the pelion and fossilbjerget formations ten sedimentary facies (table 1, facies 1–10) are identified in the pelion and fossilbjerget couplet in jameson land. the facies occur as single beds, lenses, or up to 40 m thick bed-sets. all facies were deposited in marine 820 sedimentary facies lithology, thickness of beds sedimentary structures grading and grain size lower boundary 1 coarse-grained sandstone and conglomerate. bedsets are < 1 m thick planar and trough cross-bedding or structureless normally graded or non-graded. 0.5 mm – 10 cm erosional, relief up to 0.4 m 2 coarse-grained sandstone. bedsets are < 1 m thick low-angle, planar cross-bedding or structureless non-graded. 0.5–5 mm erosional, relief up to 0.4 m 3 fineto mediumgrained sandstone. bedsets are < 2 m thick trough cross-bedding, indistinct lamination or structureless non-graded. 0.1–0.5 mm erosional, relief up to 0.3 m 4 fineto mediumgrained sandstone. bedsets are < 2 m thick trough cross-bedding, wavy bedding, cross-lamination non-graded. 0.1–0.5 mm gradational from underlying clinoform association 5 (calcareous) siltstone. beds are < 0.5 m thick structureless (100% burrowing) or planar lamination with minor burrowing non-graded. 0.05–0.25 mm erosional, wavy, relief < 10 cm 6 silty, fine-grained sandstone. bedsets are 0.3–10 m thick planar lamination, hummocky crossstratification, prod marks inversely graded or non-graded. 0.1–2 mm planar or wavy, relief < 10 cm 7 fineto coarsegrained sandstone. bedsets are 0.1–6 m thick cross-lamination, wavy-bedding, local cross-bedding inversely graded or non-graded. 0.1–2 mm erosional, wavy, relief < 6 cm 8 fineto mediumgrained sandstone. bedsets are < 6 m thick horizontal to lowangle inclined bedding, swaley cross-stratification inversely graded or non-graded. 0.1–0.5 mm gradational 9 fineto coarsegrained sandstone and conglomerate. bedsets are up to 35 m thick cross-lamination, cross-bedding, wavy-bedding or structureless non-graded. 0.1–10 cm structurally and texturally gradational or erosional 10 siltstones and finegrained sandstone. depositional units are up to 10 m thick planar lamination, hummocky crossstratification or structureless inversely graded or non-graded. 0.005–0.25 mm non-erosional, planar table 1. facies classification environments, as indicated by body and trace fossil content, sedimentary structures and the facies associations. the facies are described and interpreted in table 1. six facies associations (a–f) are recognised. they are characterised by a systematic vertical stacking of genetically related facies. the facies changes reflect the development through time of the depositional environment and its relationship to the initial bathymetry, relative sealevel changes, sediment influx and grain-size distribution. the facies associations form genetic units and do not include any major hiatuses. they are generally bounded by regionally extensive sequence stratigraphic key surfaces (discussed in the sequence stratigraphy section). the facies associations represent two large-scale 821 upper boundary fossils and bioturbation depositional processes facies association erosional, planar or wavy fragmented belemnites, bivalves, and ammonites. non-bioturbated lag formation by marine wave-ravinement shoreface (a) estuarine (b) tidal inlet (c) erosional, planar or wavy fragmented belemnites and bivalves. vertical burrows lag formation by marine wave-ravinement shoreface (a) tidal inlet (c) sand sheet (d) erosional, planar vertical burrows lag formation by marine ravinement followed by shallow offshore wave reworking shoreface (a) sand sheet (d) erosional, planar belemnites (partly fragmented) sand sheet formation by wave reworking sand sheet (d) non-erosional, planar belemnites, bivalves, ammonites, gastropods. network-forming burrows suspension fall-out, precipitation of bio-carbonate and hardground formation clinoform (e) offshore (f) structurally and texturally gradational or erosional bivalves, belemnites. horizontal and vertical burrows suspension fall-out, traction and agitation by shoaling waves offshore (f) structurally and texturally gradational or erosional bivalves, belemnites. horizontal and vertical burrows suspension fall-out, traction and agitation by waves shoreface (a) structurally and texturally gradational or erosional bivalves, belemnites. horizontal and vertical burrows suspension fall-out from sheet-flows, agitation by waves shoreface (a) erosional, wavy, with relief < 40 cm, or planar fragmented bivalves, belemnites and ammonites. horizontal and vertical burrows traction and agitation by waves, coast-parallel currents shoreface (a) estuarine (b) tidal inlet (c) sand sheet (d) clinoform (e) non-erosional, planar or erosional with relief < 10 cm belemnites, ammonites, bivalves, gastropods, arthropods and corals. horizontal and vertical burrows suspension fall-out, minor agitation and traction by waves offshore (f) depositional environment shoreface to shallow offshore, above storm wave-base shoreface to shallow offshore, above storm wave-base shoreface to shallow offshore, above storm wave-base shallow offshore, above storm wave-base offshore, below storm wave-base offshore, close to storm wave-base shoreface, above fairweather wave-base shoreface, from above storm wave-base to foreshore shoreface to shallow offshore, from below storm wave-base to above fair-weather wave-base offshore, below or close to storm wave-base depositional environments: (1) the upper shoreface to offshore transition zone, and (2) the offshore. facies associations a–e were deposited in shoreface and shallow offshore transition zone environments: a, shoreface association; b, estuarine association; c, tidal inlet association; d, sand sheet association; e, clinoform association. the offshore association (f) was deposited in a deeper-water environment. a. shoreface association the shoreface association consists of fine-grained to pebbly sandstones and conglomerates and forms coarsening-upwards units, 5–12 m thick (figs 8–11). stacked, amalgamated shoreface units are common and may reach 40 m in thickness. the foreshore to upper shoreface, middle shoreface, and lower shoreface to offshore transition 822 0 5 10 15 20 m c c c c bo un di ng su rf ac es fa ci es as so ci at io n se di m en ta ry fa ci es tr ac e fo ss ils bo dy fo ss ils c ur re nt or ie nt at io npelion fm, locality 9, 663–683 m a. sl a a 7 2 a 7 1 9 9 7 fs mse fs mse mse clay silt f m sand pebbles c f c bo un di ng su rf ac es fa ci es as so ci at io ns se di m en ta ry fa ci es tr ac e fo ss ils bo dy fo ss ils c ur re nt or ie nt at io nspelion fm, locality 10, 771–787 m a. sl clay silt f m sand pebbles c f c fs mse mse mse mse mse f d a a d 6 9 9 9 9 7 7 7 7 5 1 15 10 5 0 m fig. 8. pebbly shoreface association (a), consisting of heavily burrowed middle and upper shoreface deposits, dominated by wave-formed trough cross-bedding. the very coarse-grained nature of the deposits suggests a nearby fluvial source. pelion formation, locality 9 (fig. 2). the association belongs to sequence p1 (fig. 6). the accompanying legend (facing page) also applies to the subsequent sedimentary logs in the paper; m a. sl, metres above sea level. fig. 9. sand sheet association (d) overlain by shoreface association (a). the shoreface association consists of 8 m of mixed waveand current-influenced middle shoreface sandstones (4–12 m), overlain by 2 m of upper shoreface swash/surf-laminated sandstones. the top is formed by amalgamated upper shoreface deposits (14–18 m). the abundance of conglomerate beds suggests a nearby fluvial source. pelion formation, locality 10 (fig. 2). the association belongs to sequence p2 (fig. 6). for legend, see facing page. 823 sedimentary structures trough cross-bedding planar cross-bedding swaley cross-stratification hummocky cross-stratification structureless current ripple cross-lamination wave ripples and wave ripple cross-lamination planar bedding planar lamination indistinct lamination wavy bedding palaeocurrents, miscellaneous foreset orientation orientation and dip of inclined master bedding bipolar wave ripple foreset orientation pebbles dolerite sill coal drapes and lenses trace fossils monocraterion tentaculatum diplocraterion habichi skolithos isp. ophiomorpha nodosa gyrochorte comosa curvolithus multiplex phoebichnus trochoides planolites vulgaris helminthopsis magna taenidium serpentinum rhizocorallium irregulare thalassinoides suevica phycodes isp. arenicolites isp. degree of bioturbation body fossils belemnite ammonite bivalve plant fragments and petrified wood bounding surfaces mse marine surface of erosion fs flooding surface facies associations shoreface estuarine tidal inlet sand sheet clinoform offshore a b c d e f c sedimentary facies 1 non-bioturbated fossiliferous, coarse-grained sandstone and conglomerate lag 2 bioturbated coarse-grained sandstone lag 3 fineto medium-grained structureless to faintly laminated or cross-bedded sandstone 4 cross-bedded, cross-laminated and wavy-bedded fineto medium-grained sandstone 5 heavily burrowed calcareous siltstone and fossilliferous laminated siltstone 6 laminated silty, fine-grained sandstone 7 fineto coarse-grained, cross-laminated sandstone 8 fineto medium-grained, low-angle inclined and horizontally bedded sandstone 9 cross-laminated, cross-bedded and structureless fineto coarse-grained sandstone and conglomerate 10 offshore facies: laminated to structureless siltstone 4 o – zone deposits are dealt with separately due to the complex vertical and lateral relationships between facies. shoreface deposits dominate the sedimentary record throughout the pelion formation. foreshore and upper shoreface foreshore and upper shoreface deposits consist of fineto coarse-grained, cross-laminated, cross-bedded and structureless sandstones, and structureless or cross-bedded conglomerates (facies 9). the deposits form units, 5–40 m thick. units thicker than 5–10 m consist of amalgamated upper shoreface deposits. marine body and trace fossils occur throughout and terrestrial deposits have not been recognised. trace fossils are dominated by vertical burrows. two types of upper shoreface deposits are recognised: barred and non-barred upper shoreface sandstones (figs 8–11). barred upper shoreface deposits vary from pebbly, poorly to well sorted mediumor fine-grained sandstones; grain-size variations between individual beds may be large. sedimentary structures include a mixture of waveand current-formed cross-lamination and crossbedding, and most sandstones show an intricate mixture of swaley cross-beds, rip-channel fills, shore-parallel trough fills, bar cross-beds, wave-surf and swash bar deposits. wave-rippled sheets occur interbedded (figs 12–15). locally, cross-bed foresets show thin mudstone and siltstone drapes, and herringbone cross-bedding with bundled build-up of foresets also occurs. the lower boundary of the foreshore and upper shoreface deposits is sharp, erosional or gradational. the erosionally-based deposits show a low-angle master bedding, dipping up to 6° towards the south or south-east. the gradationally-based deposits show a horizontal to low-angle inclined master bedding, dipping less than 1–2° in a southerly direction. palaeocurrent directions show a wide scatter, although southwards orientations dominate (fig. 16). the deposits are capped by a marine surface of erosion, commonly overlain by a thin, coarsegrained sandstone or pebble lag, interpreted as genetically belonging to the overlying depositional unit (figs 17–19). the pebble lags consist of discoid or rounded quartz pebbles, 0.5–10 cm in diameter, intraformational sandstone and mudstone clasts, fragmented belemnites and bivalves, and in some cases also subangular extraformational rock fragments. fineto medium-grained, barred upper shoreface sandstones show the highest diversity of trace fossils, including diplocraterion habichi, monocraterion tentaculatum, ophiomorpha nodosa, planolites isp., thalassinoides suevica, skolithos linearis and curvolithos multiplex of the diplocraterion habichi and curvolithos ichnocoenoses of heinberg & birkelund (1984). the uppermost part of the upper shoreface units commonly shows dense burrowing by diplocraterion habichi, 824 clay silt f m sand pebbles c f c bo un di ng su rf ac es fa ci es as so ci at io ns se di m en ta ry fa ci es tr ac e fo ss ils bo dy fo ss ils c ur re nt or ie nt at io nspelion fm, locality 6, 912–934 m a. sl fs mse mse f f d 6 6 mse mse a a f 9 7 6 9 5 2 20 15 10 5 0 m fig. 10. wave-dominated shoreface association (a). the upper part shows swaley cross-stratification, with wave-formed surfaces locally with concentrations of bivalves. the uppermost part of the upper shoreface deposits is dominated by unimodally oriented cross-bedding, formed by predominantly southwards (offshore) migrating bars. pelion formation, locality 6 (fig. 2). the association belongs to sequence p1 (fig. 6). for legend, see p. 823. 825 fig. 11. a section (15 m thick) through the shoreface association (a). as illustrated on the schematic log, lower shoreface to offshore deposits (3, 4) gradually coarsen-upwards into strongly burrowed middle shoreface deposits (2), erosionally overlain by trough crossbedded upper shoreface deposits (1). the upper bounding surface of the shoreface association is a wave-cut marine surface of erosion (dashed line, mse). it is overlain by heavily burrowed, shallow offshore deposits. pelion formation, locality 9 (fig. 2). fig. 12. middle to upper shoreface sandstones showing surf-zone laminated deposits (1), wave-rippled deposits (2), swash-bar deposits (3), and deposits representing lateral migration of runnels, or rip-channel troughs (4). pen (centre, left) is 14 cm long. pelion formation, locality 9 (fig. 2). fig. 13. medium-grained sandstones of a barred, upper shoreface, showing surfand swash-zone laminated (1), waverippled (2), and swash-bar deposits (3). the bar deposits show rippled reactivation surfaces, and are interpreted as having formed during waning storm conditions. transparent part of compass is 11 cm long. pelion formation, locality 10 (fig. 2). 826 monocraterion tentaculatum and/or ophiomorpha nodosa of the diplocraterion habichi and ophiomorpha ichnocoenoses (heinberg & birkelund 1984). body fossils are mainly preserved in the pebble beds, including fragmented belemnites, ammonites and bivalves. the trace fossil assemblage, pebble and conglomerate lenses, low-angle master bedding, with a dip direction parallel to the dominant current orientations of the ripchannel and bar deposits, and the generally complex bedding suggest a highly dynamic foreshore to upper shoreface environment (clifton 1969, 1976; davis & fox 1972; dörjes & hertweck 1975; ekdale et al. 1984; curran 1985). mudstone-draped foresets and herringbone cross-bedding with bundled build-up of foresets suggest that reversing tidal currents were active, at least locally. tidal processes seem in general to have been overprinted by wave action and associated wave-induced currents. the wave-ripple cross-laminated sediments represent aggradational fair-weather deposits, whereas cross-beds were formed by stationary and seawards (southwards) migrating bars during storms (clifton et al. 1971; davis & fox 1972; clifton 1976; hunter et al. 1979). some shoreface deposits represent preserved shore-parallel bars, with wave-ripple cross-lamination and cross-bedding, and runnels with wave-ripple laminated fill, interpreted as fair-weather deposits (figs 12–15). the conglomerate lenses and the low-angle inclined laminated beds of barred upper shorefaces are interpreted as bar and swash-/surf-zone deposits (clifton fig. 14. foreshore deposits with parallel to low-angle, inclined, fineto mediumgrained sandstones of fair-weather origin, erosionally overlain by coarse-grained, belemnite-rich, pebbly storm sandstones. bioturbation by ophiomorpha nodosa. pen (lower, right) is 14 cm long. pelion formation, locality 9 (fig. 2). fig. 15. sandstones showing horizontal to low-angle inclined bedding (1), cut by south to south-east oriented large-scale rip-channel trough fill (2), showing epsilon-infill pattern, caused by eastwards lateral migration of the rip-channel. 3, wave-rippled beds. pelion formation, locality 9 (fig. 2). total thickness of (2) is c. 2 m. 827 1969, 1976; hunter et al. 1979). the presence of conglomerates indicates proximity to a fluvial source, and they are interpreted to have been deposited in wavereworked fluvial mouth bars and adjacent beaches. the sediments are thought to have been delivered to the shoreface regions through braided fluvial channels, which must have been shallow, because no channel deposits survived subsequent transgressive marine erosion. the seawards inclination of the master bedding in some upper shoreface sandstones is interpreted to indicate that shoreface progradation occurred during high energy conditions with high sediment influx (cant 1991; swift & thorne 1991). the master bedding marks an abrupt upwards change from lower to middle shoreface aggradation to dominantly upper shoreface progradation. the erosional surface which in many cases separates middle or lower from upper shoreface deposits fig. 17. a fine pebble lag, 1 m thick, bounded below by a marine surface of erosion (mse) and topped by a flooding surface (fs). the lag apparently forms a single cross-bed with foresets dipping up to 20° towards the south, indicated by the fracture pattern. the formation of the lower boundary (mse) was followed by colonisation of diplocraterion habichi and monocraterion tentaculatum during an early stage of water deepening (arrow 1). during the subsequent winnowing and deposition of coarse material, the burrowing organisms still inhabited the sediment (arrow 2). the deposits become increasingly bioturbated upwards with nearly 100% burrowing at the top, and d. habichi descending from the top flooding surface. pelion formation, locality 7 (fig. 2). current-ripple orientations in middle shoreface deposits, localities 10 and 11 wave-ripple crestline orientations in middle shoreface deposits, locality 11 current-ripple orientations in upper shoreface deposits, localities 6, 6a, 7, 9, 9a, 10, 11, 14 wave-ripple crestline orientations in upper shoreface deposits, localities 6, 6a, 7, 9, 9a, 10, 11, 14 current-ripple orientations in sand sheet deposits, localities 7, 8, 10, 11 wave-ripple crestline orientations in shallow offshore sand sheet deposits, localities 7, 8, 10, 11 equal area circle = 11% n = 54 equal area circle = 14% n = 25 equal area circle = 13% n = 47 equal area circle = 13% n = 28 equal area circle = 13% n = 32 equal area circle = 8% n = 25 fig. 16. current-ripple foreset azimuths and wave-ripple crestline orientations in shoreface and shallow offshore sandstones. data from several localities are included. the shoreface and shallow offshore units were most likely deposited by, or in connection with, progradation of roughly east–west-oriented shorelines. each shoreface unit prograded from the north. this type of data does not allow precise interpretation of the shoreline orientation, due to possible wave and current interference patterns in the nearshore areas. the data may, however, indicate a general trend, as they show a pronounced lateral persistence. 828 fig. 18. non-winnowed deposits, interpreted to have formed during transgressive deepening. they overlie a marine surface of erosion (mse), and consist of massive or weakly laminated silty, fine-grained sandstones (arrow 1) overlain by coarse-grained, even, parallel bedded and cross-bedded sandstones (arrow 2). bioturbation by diplocraterion habichi and monocraterion tentaculatum is concentrated in the uppermost part, descending from the top flooding surface (fs). the coarse grain size of the upper bed suggests that it was deposited as a shoreface-connected sand-sheet during shoreface retreat. pelion formation, locality 7 (fig. 2). fig. 19. carbonate-cemented pebble lag (uppermost), interpreted to have formed during transgressive erosion. the lag overlies upper shoreface to foreshore, cross-bedded sandstones with an erosional base, representing a marine surface of erosion (mse, dashed line). the upper boundary is a marine flooding surface (fs), which is overlain by laminated, offshore siltstones. the conglomerate bed can be traced laterally over an area of 4 x 2 km. hammer shaft for scale, 4 cm across end. pelion formation, locality 10 (fig. 2). fig. 20. sandstones showing horizontal to low-angle inclined beds (1), overlain by swaley cross-stratified beds (2), cut by s–se-oriented rip-channel trough fills (3). wave-ripple cross-laminated beds occur at different levels (4). parallel-laminated surf-zone beds are also present (5). the sandstones are interpreted as wavedominated shoreface deposits. hammer head (encircled) is 18 cm long. pelion formation, locality 7 (fig. 2). was formed by storm-wave and current scour in front of the prograding upper shoreface. the surface is diachronous, and it has no regional significance in correlation (nummedal & swift 1987; swift et al. 1991). the southwards orientation of the inclined upper shoreface master bedding and the dominant east–west orientation of wave-ripple crests in both lower and upper shoreface deposits indicate that shorelines prograded southwards (fig. 16). the overall grain-size gradients suggest sediment influx from the north-west. non-barred, wave-dominated upper shoreface sandstones were deposited in some areas, especially at localities 6, 6a, and 7 (fig. 20). the sandstones are generally medium-grained, well sorted, and dominated by swaley and wave-ripple cross-laminated beds, erosionally cut by north–south oriented rip-channels and gutter casts (facies 9). the deposits always gradationally overlie fineto medium-grained, low-angle inclined and horizontally-bedded sandstones of facies 8 (fig. 20). the deposits are capped by a marine surface of erosion, overlain by erosionally-based swaley cross-beds. the degree of bioturbation is generally low, and planolites vulgaris and skolithos isp. have been identified. belemnites occur scattered throughout the facies. the dominance of swaley cross-bedding and waveripple cross-lamination, erosionally cut by rip-channels and gutter casts, suggests deposition in a storm-wave dominated upper shoreface (plint 1988; bhattacharya & walker 1991). the fossil assemblage indicates marine conditions. the good sorting suggests a position well away from fluvial input sources. the predominant north–south orientation of gutter casts and rip-channels suggests an east–west oriented shoreline. the generally low degree of bioturbation may reflect the non-barred character of the shoreline, lacking protected areas. middle shoreface the middle shoreface deposits consist of 1–10 m of fineto locally coarse-grained sandstones (facies 7, 8), containing scattered bivalves and belemnites. waveformed ripples and mega-ripples characterise most of the deposits, including lenses of cross-bedded, mediumto coarse-grained sandstones, and erosionally-based fineto coarse-grained, laterally aggradational trough fill sandstones, up to 0.5 m thick (facies 7). bioturbation is generally very intense. horizontal traces dominate in the wave-rippled beds, and dominant trace fossils are: taenidium serpentinum, gyrochorte comosa, planolites isp., phycodes isp., and thalassinoides suevica, whereas vertical traces dominate in the cross-beds, including notably monocraterion tentaculatum, and skolithos linearis. the trace fossils are referred to the planolites, curvolithus and diplocraterion habichi ichnocoenoses of heinberg & birkelund (1984). the deposits may form bedsets, up to 8 m thick, of fineto medium-grained, low-angle inclined to flat-bedded sandstones, interbedded with wave-rippled beds and overlain by swaley cross-bedded sandstones of the upper shoreface (facies 8; figs 10, 20). palaeocurrent directions are towards the west, through south to the east, but south to southeast orientations dominate (fig. 16). the lower boundary of the deposits is erosional and wavy, or there is a gradual upwards change from hummocky cross-stratified or even, parallel-laminated beds, and the boundary is placed at the base of the wave-ripple laminated beds. the deposits may directly overlie marine lags without any intervening offshore or lower shoreface sediments. the upper boundary of middle shoreface deposits is in most cases erosional and sharp, showing a relief of up to 0.4 m, or there may be a textural and structural gradational change into upper shoreface deposits, showing both horizontal traces, and vertical traces. the dominant fineto medium-grained, ripple-laminated beds were deposited under the influence of fairweather waves. the dominance of horizontal traces indicates overall slow deposition, whereas the change to vertical traces represents a seawards translated upper shoreface fauna (dörjes & hertweck 1975; frey 1975; ekdale et al. 1984; pemberton & frey 1985). fair-weather deposition was punctuated by storm erosion, reworking and deposition, represented by the mediumto coarse-grained, cross-bedded lenses. the cross-beds are interpreted to have been deposited in rip-channels, on rip-channel mouth-bars and as mega-ripples, formed during seawards transport of sediments during major storms. the current orientations obtained from the ripchannel and associated mouth-bar deposits, and waveripple crestline orientations suggest an east–west orientation of the coastline (fig. 16). the large spectrum of preserved waveand current-formed sedimentary structures suggests that deposition was influenced both by storm and fair-weather wave processes, and that the water depth was in the range of 5–15 m during deposition. similar deposits have been described by clifton (1976), hunter et al. (1979), dabrio & polo (1981), and swift et al. (1991). the flat-bedded sandstones are interpreted as sand sheets deposited from density currents, fed by laterally 829 830 extensive erosion of upper shoreface deposits during storms. this process has been well described by howard & nelson (1982) and nelson (1982). the flat-bedded sandstones may also have formed rip-channel mouth bars, deposited from a mixture of bedload and suspension fall-out, as indicated by the associated ripchannel fills which occur locally. channelised and non-channelised sediment transport probably took place simultaneously during storms; this is indicated by the small size of the preserved rip-channels, compared to the thickness and lateral extent of the flat-bedded sandstones (fig. 20). shoreface associations may lack lower shoreface and offshore transition zone deposits, and consist of middle shoreface deposits in the lower part, overlying lags or marine regressive surfaces of erosion (as shown in figs 8, 9). this is thought to indicate that the shoreface prograded into very shallow water, as described for cretaceous deposits of the us western interior seaway by hart & plint (1993). lower shoreface to offshore transition zone the lower shoreface to offshore transition zone deposits are characterised by a backstepping, aggradational to forestepping facies stacking pattern. the deposits have a basal lag, ranging from a layer only one pebble thick up to 1 m thick (facies 1, 2), or rest directly on a strongly burrowed marine erosion surface, locally overlain by less than 2 m of fineto medium-grained, structureless, faintly-laminated or cross-bedded sandstone (facies 3; figs 17–19). the basal deposits are overlain by less than 0.5 m of strongly burrowed calcareous siltstone or laminated siltstone (facies 5; figs 11, 21), locally rich in bivalves, belemnites and ammonites, sharply overlain by laminated, silty, fine-grained sandstone, 0.5–15 m thick (facies 6). the latter deposits consist of intercalated even, parallel-laminated or hummocky cross-stratified beds, or show an upwards transition from horizontally laminated to hummocky cross-stratified beds. wave ripple-formsets locally occur in the uppermost part of the deposits. burrows of gyrochorte comosa, taenidium serpentinum, planolites vulgaris and helminthopsis magna (planolites ichnocoenosis) are restricted to the even, parallel laminated beds. marine body fossils are found locally throughout. the association is capped by a sharp, wavy boundary, or shows a gradational contact to overlying fineto coarse-grained pebbly, rippled or cross-bedded sandstones of the middle or upper shoreface (figs 8–11). the association was deposited in the lower shoreface to offshore transition zone. the coarse-grained basal lags and overlying calcareous deposits are interpreted as having been formed during and after transgressive erosion of the underlying shoreface deposits, and indicate an increase in water depth from above fair-weather wave base to near or below storm wave base (nummedal & swift 1987; nummedal et al. 1993; swift et al. 1991). the lags were formed during a time of winnowing or erosion with no net deposition, and the overlying condensed facies (facies 5, 6) were deposited slowly from suspension fall-out, occasionally influenced by weak currents. the hummocky cross-stratified beds fig. 21. a calcareous siltstone bed with a thalassinoides suevica polygonal network (seen from below). the bed is 0.3 m thick; it is interpreted as representing a firmground, developed during a time interval of reduced deposition, during a rise in relative sea level (cf. bromley 1975). pen (encircled) is 14 cm long. pelion formation, locality 11 (fig. 2). were formed by storm wave-oscillation, and individual beds represent single storm events (harms 1975; brenchley et al. 1986, 1993). deposition took place at a minimum water depth of 5–15 m, interpreted from the thickness of the overlying middle and upper shoreface deposits, and the preserved bedforms and trace fossils. the marine trace fossils of the horizontally laminated beds are typical of an environment dominated by slow suspension fall-out (seilacher 1967; frey 1975; heinberg & birkelund 1984). b. estuarine association preserved estuarine deposits are rare, and have only been found at two localities. they are described below and include channel and shoal deposits. estuarine channel estuarine deposits are exposed in the valley between the mountains of pelion and parnas, where estuarine deposits of the pelion formation erosionally overlie the sortehat formation (fig. 2, locality 10; see also figs 22, 23). erosional incision reaches a presumed maximum of 30–40 m in the central part of the exposure, which is 1.5 km long and is oriented approximately east–west, perpendicular to the axis of the incised valley, as deduced from current orientations and large-scale bed geometries. the estuarine deposits are 30–40 m thick and are dominated by medium-grained sandstones with a few conglomerate beds (facies 9). two channel sandstone bodies occur, separated by thin spit deposits. the upper estuarine channel unit is overlain by strongly burrowed offshore sandstones and siltstones of facies 6 and 7 (fig. 22). the basal erosion surface is draped by a discontinuous, allochthonous, vitrinite coal layer, up to 4 cm thick, overlain by a fine-grained sandstone, up to 0.5 m thick, rich in coal debris and mica (fig. 23). the over831 bo un di ng su rf ac es fa ci es as so ci at io ns se di m en ta ry fa ci es tr ac e fo ss ils bo dy fo ss ils c ur re nt or ie nt at io ns pelion fm, locality 10, 550–595 m a. sl clay silt f m sand pebbles c f c mse f mse fs fs b a 6 9 9 9 9 7 7 7 fs fs mse b f 9 7 9 7 7 m 45 40 35 30 25 20 15 10 5 0 c c c c c c c sortehat fm * fig. 22. estuarine channel association (b). the inclined master bedding of the lower channel unit is indicated (3.5–29 m). the lower channel unit is capped by spit sandstones (29–32.5 m), which are overlain by the upper channel unit (32.5–41 m). bioturbated, fully marine offshore sandstones occur on top of the upper channel-fill. asterisk indicates inferred floodplain deposits. pelion formation, locality 10 (fig. 2). for legend, see p. 823. 832 lying sandstones are characterised by low-angle master bedding, dipping c. 4° towards the east, representing the fill of a major, laterally migrating estuarine channel (morton & mcgowen 1980; cotter 1983; yang & nio 1989; sha & de boer 1991). the general current orientation was towards the south, but other directions are also observed. large-scale cross-beds, up to 2 m thick, dominate, showing planar and tangential foresets dipping towards the south. the concentration of organic debris is commonly high along bottomsets, and may form allochthonous coal beds up to a few centimetres thick. petrified wood fragments are abundant throughout. skolithos linearis and diplocraterion habichi occur at eight levels from 4 m above the base of the deposits and upwards. the cross-beds may be separated by cross-laminated sandstone beds or bedsets 0.2–3 m thick, locally containing coal debris and centimetrescale, discoidal quartz pebbles. ripple crests are generally oriented east–west. the lower channel unit is overlain by three coarsening-upwards units, each c. 1 m thick, with a horizontally bedded or cross-laminated lower part and a planar or trough cross-bedded upper part (fig. 22). the top unit is erosionally overlain by the upper channel fill unit, which is up to 10 m thick, consisting of fairly well sorted, medium-grained sandstones, except for the lowermost metre, which is coarse-grained. the channel fill shows 1–20 m wide troughs, which cut into each other, with margins typically inclined less than 10°. the trough fills show concordant bedding and lamination, and occasionally wave-rippled surfaces, but lateral accretion surfaces also occur. palaeocurrent orientations are consistently towards the south. the channel deposits are overlain by an erosionally-based sandstone unit, 3 m thick, rich in coal debris, larger fragments of plants, and petrified wood. the sandstones show wave-ripple cross-lamination, and locally wave-ripple formsets. this upper unit shows moderate to very strong burrowing, including diplocraterion habichi, monocraterion tentaculatum, skolithos linearis and teichichnus isp. the succession is interpreted as the fill of a minor incised valley, cut into the dark offshore mudstones of the sortehat formation by fluvial erosion during a time of sea-level lowstand. the vitrinite coal layers and finegrained sandstones rich in coal debris of the basal valley-fill deposits are interpreted as the erosional remains of floodplain sediments, which escaped transgressive erosion during the transformation of the river mouth into an estuary. the large-scale cross-beds of the lower estuarine channel unit were formed by migration of sub-channels and bars on the inclined accretion surface of the channel margin. the channel sandstones are interpreted as freshwater-influenced ebb-tidal deposits, based on the large amount of plant material, the dominance of southwards migrating bed-forms, and the restriction of marine trace fossils to channel inactivity surfaces (barwis 1985; yang & nio 1989; allen 1991). the fine-grained wave-rippled beds and bedsets rich in coal debris and with scattered discoidal pebbles that occur at different levels on the inclined master bedding surface, were possibly formed by wave reworking during channel inactivity. the predominant east–west orientation of wave-ripple crests is interpreted to reflect the north–south fig. 23. basal 4 m of the estuarine channel sandstones, filling a minor valley incised into the sortehat formation. the deposits belong to sequence p1 (fig. 6). sb1, lower sequence boundary of the pelion formation, separating estuarine deposits from underlying marine silty shales of the sortehat formation; 1, floodplain deposits, representing the oldest deposits in the incised river valley; 2, allochthonous vitrinite and durite coal layer; 3, base of lower estuarine channel unit (dotted line); 4, estuarine channel sandstones. hammer for scale (at sb1). pelion formation, locality 10 (fig. 2). view is towards the south-west. orientation of the estuary. the three small-scale coarsening-upwards units, overlying the estuarine tidal channel deposits are interpreted as spit deposits. they were formed by laterally migrating spits, which filled the accommodation space left above the channel fill after the active channel had moved laterally by erosion and accretion (terwindt 1971; yang & nio 1989; allen 1991). the upper channel unit is also interpreted as a freshwater-influenced ebb-tidal estuarine channel fill, based on the invariably southwards-oriented, erosionallybased large to giant-scale foresets, and the lack of marine fossils and trace fossils (dalrymple et al. 1992). the amalgamated character of the channel deposits indicates a high sediment influx to the estuary, larger than the rate of creation of new accommodation space. the uppermost 3 m of burrowed, wavy bedded sandstones are interpreted as a sand sheet, deposited during marine flooding of the estuary (campbell & oaks 1973; dalrymple et al. 1992). estuarine shoal estuarine shoal deposits are exposed low on the south slope of the olympen mountain (fig. 2, locality 8). the exposed part of the unit is 14 m thick and consists of fineto medium-grained sandstones, with a few conglomeratic sheets and lenses, each less than 10 cm thick. clay and coaly detritus form thin drapes and layers. the deposits overlie and are overlain by shallow offshore sediments; the lower boundary is only exposed at a single place, where it is overlain by bottomsets of a planar cross-bed, showing bundled, low-angle foresets, rich in coal debris. the sedimentary structures are dominated by large-scale trough and planar cross-bedding, low-angle trough cross-bedding, wavy bedding, and planar-bedding (facies 7, 9). beds are generally less than 0.5 m thick. in the lowermost part of the unit, individual crossbeds can be traced in a southwards, down-current direction for more than 80 m without significant variation in thickness. they show bundle-wise up-building of sigmoidal, tangential or angular foresets, which may be draped by coaly debris and dark clay. they are separated by wavy bedded, fine-grained sandstones. the middle part of the unit is dominated by interbedded low-angle, cross-bedded sandstone with shallow crossbedded sandstone trough fills, up to 1 m wide, with foresets draped by coal debris, and swaley cross-stratification. cross-beds show foreset azimuths from 90° through 180° to 220°. bed boundaries of the lower and middle part of the unit mostly have a wavy appearance, and can in most cases be followed throughout the exposure. the upper part is dominated by non-bioturbated, swaley cross-stratified sandstones. ophiomorpha nodosa burrows occur in the lower and middle part of the unit, where they descend in great numbers from bedding planes. the estuarine shoal interpretation is based on the internal sedimentary structures and the stratigraphic position, sandwiched between offshore sandstones and siltstones. the presence of sandstones rich in coal debris, coal layers at the base of troughs, small-scale waveformed ripples, and the dominance of ophiomorpha nodosa indicate deposition at very shallow water depths, not far from a terrestrial source (clifton 1976, 1983; barwis 1985; curran 1985; pollard et al. 1993). the basal estuarine shoal beds with bundled foresets, are interpreted to have formed by migration of subto intertidal sand waves, under the influence of neap–spring fluctuations of tidal currents. the wave-rippled beds are interpreted as slack-water sediments. these lower deposits are interpreted to have accumulated in shallow tidal channels, probably less than a few metres deep, based on the thickness of the channel fills. the middle part of the estuarine shoal deposits strongly resembles modern tidal shoal sediments (de raaf & boersma 1971; terwindt 1971; reineck 1972; boersma & terwindt 1981). the non-bioturbated upper part of the deposits is interpreted to reflect a general increase in marine storm and wave influence. the vertical facies development of the estuarine shoal deposits closely resembles tidal flat associations from the lower cretaceous fall river formation, wyoming (campbell & oaks 1973) and modern examples (de raaf & boersma 1971; clifton 1983; terwindt 1988). the estuarine shoal deposits are thought to represent non-erosional drowning of an estuarine shoal area (campbell & oaks 1973; dalrymple et al. 1992). the top marine surface of erosion is interpreted to have formed during the final marine transgression of the estuary. c.tidal inlet association tidal inlet deposits are exposed at two stratigraphic levels at trefjord bjerg in the easternmost part of the exposure belt (fig. 2, localities 6, 6a, 7). the lower and upper inlet units are 20–25 m and 20–35 m thick, respectively. both units can be traced laterally east–west for 1.5 km. the lower can be followed north–south for 600 m, and the upper for several kilometres, without 833 any change in facies pattern. locally, mouth-bar deposits are present in the lower part of the association. in the absence of basal mouth-bar deposits, the tidal inlet units erosionally overlie offshore siltstones and sandstones. the mouth-bar deposits are up to 8 m thick, and consist of fineto medium-grained currentand wave-ripple cross-laminated sandstones (facies 7). the upper part of the mouth-bar deposits is especially rich in coal debris, plant fragments, bivalves and belemnites. the mouth-bar deposits are erosionally truncated by the base of the tidal inlet channel fill, and locally a large number of diplocraterion habichi burrows descend from the erosion surface. the inlet channel fill consists of mediumto coarse-grained, generally well-sorted sandstones (facies 9), with very subordinate conglomerates. vertical trends in grain size have not been recorded (figs 24, 25). the master bedding of the tidal inlet sandstones is inclined 4° in an easterly direction. the master bedding surfaces show tangential to angular lower bedding contacts and are erosionally truncated updip by a marine erosion surface. the inclined strata are 0.2–2 m thick and show planar and trough cross-bedding, massive bedding or ripple cross-lamination. wave-ripples may occur at bed boundaries. reactivation surfaces are common, both at bed and master bedding scale. herringbone cross-bedding is common. southerly palaeocurrent directions dominate with subordinate northerly directions, perpendicular, or at a very high angle, to the dip direction of the master bedding surfaces. scattered diplocraterion parallelum, planolites isp., phycodes isp. and a single taenidium serpentinum were observed in the channel deposits. diplocraterion habichi and monocraterion tentaculatum descend in large numbers from the marine erosion surface at the top of the inlet channel sandstones. ammonites occur in the lower part of the association, whereas belemnites and bivalves are found scattered in the upper part. 834 bo un di ng su rf ac es fa ci es as so ci at io ns se di m en ta ry fa ci es tr ac e fo ss ils bo dy fo ss ils c ur re nt or ie nt at io nspelion fm, locality 7, 858–895 m a. sl clay silt f m sand pebbles c f c fs mse f d 6 5 1 mse c 7 f 6 9 30 25 20 15 10 5 0 m fig. 24. tidal inlet association (c), consisting of large-scale planar cross-bedded, mediumto coarse-grained sandstones (10–31 m), erosionally overlying finer grained, cross-laminated, and cross-bedded inlet mouth deposits (3.5–10 m). the largescale bedforms in the inlet deposits indicate that powerful currents were active at the time of deposition. bedforms mainly migrated towards the south, perpendicular to the azimuth of the master bedding planes. ammonites found in the offshore deposits below the inlet channel (at 0.5 m) belong to the cranocephalites borealis β faunal horizon. pelion formation, locality 7 (fig. 2). for legend, see p. 823. 835 the body and trace fossils indicate that normal marine salinities prevailed, at least periodically. the restriction of trace fossils to distinct levels suggests that rates of deposition and erosion of the active depositional surfaces were generally too high for colonisation, which only occurred during longer periods of bedform stillstand. the tidal inlet interpretation is based on the lateral association with shoreface deposits, and by the internal sedimentary structures, including the inclined master bedding (terwindt 1971, 1981; nio & yang 1991). the eastwards inclined master bedding reflects lateral migration of the inlet, and the presence of up to 2 m thick cross-beds suggests deposition from powerful tidal currents. the common occurrence of herringbone cross-bedding suggests that both ebband flood-currents operated in the channels, although ebb-oriented currents dominated, as deduced from the dominant offshore, southwards palaeocurrent direction. the generally well-sorted nature of the deposits suggests that sorting processes operated so as to maintain a uniform grain size in the channels; finer grained sediment bypassed the channels to be deposited in adjacent shoreface, channel-mouth, and offshore environments. the sediment sources for the channel fills were presumably mainly cannibalised shoreface deposits, and to a lesser extent, primary fluvial sediments. the tidal inlets were subtidal as the inclined master bedding covers the full thickness of the preserved deposits. the depositional environment was probably somewhat similar to recent tidal inlets described from the dutch north sea coast by yang & nio (1989) and sha & de boer (1991). the bases of the inlet channels form channel diastems, and are locally strongly incised. they may be considered as unconformities as discussed below in the sequence stratigraphy section. the tidal range has not been calculated due to lack of relevant data, but is estimated to have been in the range of 1–2 m, by comparison with bedform patterns of the modern north sea. the shoreface associations of the pelion formation show little evidence for tidal inlets, which would have dominated if the tidal range was larger. the modern example described by sha & de boer (1991) shows that at a tidal range of 1–2 m, the tidal inlets and estuaries are the only part of the sea-facing environment that are dominated by tidal processes. adjacent shorefaces are dominated by waves, long-shore currents and shorenormal rip-currents which are not systematically reversing. this is compatible with the data from shoreface and tidal inlet channel deposits of the pelion formation. d. sand sheet association cross-bedded sandstone sheets, 0.5–5 m thick, which are erosionally based and have sharp upper boundaries, are found in central and northern jameson land. they consist of fineto coarse-grained quartzose sandfig. 25. upper tidal inlet association (c) from the lower part of the section in fig. 24. the master bedding (mb, dashed) forms giant-scale foresets or clinoforms, internally showing parallel bedding (1), large-scale planar cross-bedding (2), and homogeneous, structureless bedding (3). 4, wave-ripple cross-bedding. person for scale (170 cm). pelion formation, locality 7 (fig. 2). stones. the sorting is locally very good, making foresets almost invisible. belemnites, ammonites and bivalves occur scattered. individual sand sheets are tabular, and have in some cases been traced over more than 9 km2. the sheets may be solitary, less than 1.5 m thick and interbedded with shallow offshore to middle shoreface deposits, or form vertically stacked successions, in which the individual sheets are separated by thin offshore or lower shoreface units. some thin sheets consist of a single bed with planar cross-bedding. the upper boundary is a sharp, planar surface, or is marked by well-preserved wave-formed megaripples. sedimentary structures in the association include trough and planar cross-bedding, cross-lamination and even, parallel lamination; the structures do not occur in any clear vertical succession (figs 9, 26). herringbone cross-bedding, showing bundled up-building of foresets, is locally observed. internal scour fills showing lateral accretion are present in some units, reflecting local, confined currents. a wide range of palaeocurrent directions are observed in individual sand sheets, but southwards directions predominate (fig. 16). the sand sheet units are generally pervasively burrowed by taenidium serpentinum, planolites isp. and gyrochorte comosa representing the curvolithos ichnocoenosis. diplocraterion habichi, monocraterion tentaculatum and skolithos linearis or ophiomorpha nodosa of the diplocraterion habichi and ophiomorpha ichnocoenoses commonly descend in large numbers from the upper bounding surface, but may also occur within the units. thalassinoides suevica and phoebichnus trochoides may also occur in large numbers. in jameson land, phoebichnus trochoides has previously been described from fine-grained offshore deposits of the fossilbjerget formation (bromley & asgaard 1972; heinberg & birkelund 1984), but also occurs in pelion formation sandstones in wollaston forland (surlyk & clemmensen 1983). the body and trace fossils of the deposits indicate a fully marine environment. the sand sheet units resemble upper shoreface sandstones, but differ in many aspects (surlyk & noe-nygaard 1991). the occurrence 836 bo un di ng su rf ac es fa ci es as so ci at io ns se di m en ta ry fa ci es tr ac e fo ss ils bo dy fo ss ils c ur re nt or ie nt at io nspelion fm, locality 8, 577–610 m a. sl clay silt f m sand pebbles c f c mse f 6 mse mse mse mse d d d d f mse d mse f f f f mse mse mse mse fs d d d 9 7 6 7 9 9 7 7 9 9 9 9 6 6 6 9 m 35 30 25 20 15 10 5 0 fig. 26. intercalated sand sheets (association d) and laminated transitional offshore sandstones (association f). the presence of pebbles and coarse-grained sandstone suggests that the sand sheets were closely associated with a shoreface during deposition. pelion formation, locality 8 (fig. 2). for legend, see p. 823. of thalassinoides suevica and phoebichnus trochoides is a distinct feature, which only characterises sand sheet and more fine-grained offshore deposits. the close association with lower shoreface to offshore fine-grained silty sandstones is also a distinct feature of the association. the wave-formed, mega-ripple formsets at the top of the sheets were formed by storm waves during abandonment. the trace fossil distribution shows that deposition was relatively slow with minor wave-reworking, alternating with times of storm-wave and current erosion and deposition. the horizontal traces of planolites isp., gyrochorte comosa and taenidium serpentinum are interpreted to have been made during slow fairweather deposition, whereas diplocraterion habichi, monocraterion tentaculatum, skolithos linearis and ophiomorpha nodosa probably represent a seawards translated upper shoreface assemblage (dörjes & hertweck 1975). phoebichnus trochoides and thalassinoides suevica, which occur in the top of the sand sheets, were formed after abandonment, and mark the return to quiet offshore conditions, with slow deposition from suspension fall-out (bromley & asgaard 1972; surlyk & clemmensen 1983; heinberg & birkelund 1984). the occurrence of coarse-grained sandstone in some of the thick sand sheet units suggests a nearshore position, and some sheets pass into progradational shoreline deposits towards the north-west or north. the presence of herringbone cross-bedding probably indicates reversing tidal currents. the co-existence of smalland large-scale symmetrical and slightly asymmetrical ripples indicates that the depositional surface was influenced by both fair-weather and storm waves and currents and that deposition took place above fair-weather wave base. e. clinoform association the clinoform association is only found at localities 3 and 4 (fig. 2), where it forms the ugleelv member of the pelion formation (surlyk 2003, this volume, fig. 5). it consists of a single tabular, clinoform bed or giantscale cross-bed, 8–35 m thick (locality 3), or two stacked beds, the lower up to 25 m thick and the upper up to 8 m thick (locality 4; figs 27, 28). the upper clinoform bed at locality 4 wedges out towards the south, and is not present at locality 3 (fig. 27). the association consists of well sorted fineto coarse-grained sandstones, with tangential clinoforms dipping up to 10° towards the ese. systematic lateral grain-size variations are not observed. the clinoform beds have been traced over c. 50 km2. towards the south and east, they rapidly give way to offshore siltstones and thin, fine-grained sandstones of the fossilbjerget formation. the association has not been traced west or north of locality 4, due to lack of exposure. the clinoform beds have non-erosional bases, and clinoforms pass downwards into bottomsets composed of strongly burrowed, structureless or horizontally bedded fine-grained silty, carbonaceous sandstone (fig. 29). the clinoform beds are flat-topped, bounded by a sharp erosional surface (figs 27, 28, 30). erosion probably removed less than a few metres of the units, as deduced from the thickness of locally preserved wave-rippled and swaley cross-bedded topset deposits. the upper decimetres of the clinoform beds are strongly carbonate cemented. diplocraterion habichi and monocraterion tentaculatum descend in large numbers from the upper surface, which is overlain by strongly burrowed calcareous siltstones of the fossilbjerget formation (facies 5). reactivation and inactivity surfaces are developed throughout the clinoform beds, and are reflected by two types of bundling of the clinothems. stillstand-surfaces are overlain by downlapping low-angle clinothems and are characterised by strongly burrowed carbonaceous deposits rich in fossils and plant fragments. firmground-type burrowing of thalassinoides isp. and other trace fossils dominates (fig. 31). the reactivation surfaces are characterised by erosional truncation of low-angle clinoforms by higher-angle clinoforms. intrasets are 0.2–1 m thick, mainly planar cross-beds with sharp boundaries, locally showing small-scale scours. some clinothems consist of a cross-bedded intraset or stacked cross-laminated, cross-bedded, planar bedded, or structureless sets. the intrasets locally show bundling of foresets, a lateral development from tangential to sigmoidal foresets, single and double clay drapes of bottomsets and foresets, reactivation surfaces, and herringbone cross-stratification. wave-eroded intraset boundaries are common. the foreset azimuths of intrasets are generally towards the south, and the intrasetforming ripples migrated perpendicular to the dip or obliquely down the ese-dipping clinoforms. ammonites, belemnites and bivalves occur scattered throughout the association, and are present in large numbers on the top surface of the clinothems. crinoid ossicles are abundant in bottomsets of the distal, wedging-out parts of the clinothems. petrified wood fragments and carbonised plant material are common. trace fossils include diplocraterion habichi, monocraterion tentaculatum, skolithos linearis, thalassinoides suevica, curvolithos multiplex, planolites vulgaris, teichichnus 837 838 m 20 15 10 5 0 m 30 m 30 m 40 m 35 20 15 10 5 0m 30 25 20 15 10 5 0 35 30 25 20 15 10 5 0 25 20 15 10 5 0 25 20 15 10 5 0 lo c. 6 a lo c. 4 lo c. 4 lo c. 3 lo c. 3 lo c. 2 n or th 35 k m 6. 5 km 2. 5 km 2 km 6 km so ut h 5 1b 3 3 4 4 4 3 2 4 4 3 2 9a 1b c ls l g r sd c ls l g r sd c ls l g r sd c ls l g r sd c ls l g r sd c ls l g r sd m ud st on e si lts to ne sa nd st on e pe bb le s fa un al h or iz on s u pp er cl in of or m un it lo w er cl in of or m un it 8° 8° 8° 7° 8– 10 ° isp., and phycodes isp. of the cochlichnus and diplocraterion habichi ichnocoenoses of heinberg & birkelund (1984). the clinoform beds were formed by progradation of thick marine sandbodies towards the ese. the coarse grain size suggests that deposition took place in connection with shoreface progradation, and that powerful currents, which probably periodically reached 1–1.5 m/s (indicated by the presence of traction-transported coarse sand), shaped the clinoform beds. the fossil assemblage indicates a fully marine environment. the depositional environment was tidally-influenced and shallow marine. the coexistence of smalland large839 facing page: fig. 27. sections measured through the clinoform complex (association e) of the ugleelv member, exposed at localities 3 and 4 (fig. 2), extending for 11 km in a north–south direction. the ammonites found just below the clinoform complex belong to the cranocephalites borealis β faunal horizon (1b); the ammonite faunal horizons (1–21) shown on this and subsequent figures are listed in full on fig. 37. a cranocephalites indistinctus fauna is found in the basal beds of the lower clinoform unit (2), a cranocephalites intermissus fauna (3) is found high in the unit, and a cranocephalites gracilis fauna (4) is found on top of the lower unit at locality 4, and on the top of the upper unit throughout the area. biostratigraphical correlation to locality 6 (to the north) and locality 2 (to the south) is indicated. to the north, at locality 6a, cranocephalites carlsbergensis (5) occurs immediately above the cranocephalites borealis β fauna. 9a indicates the arctocephalites delicatus faunal horizon. for full legend, see p. 823. fig. 28. the two clinoform units (association e) of the ugleelv member (fig. 27). the view is towards the south. the two units are indicated as i and ii. the base of the lower unit (i) is shown by the lowermost dashed line to the right. the top of the lower unit is dashed and is marked tse. clinoforms in the lower unit are indicated by the intervening dashed lines. the top of unit ii is marked tse and forms the top surface in fig. 27. pe fm, uppermost pelion formation; fb fm, basal part of fossilbjerget formation. locality 4 (fig. 2). fig. 29. bottomsets of the lower clinoform unit (fig. 27). compass for scale (arrow). the view is towards the east. a tidal megaripple is seen below the compass; it shows bundled foresets, with azimuths towards the south. bioturbated sandstones occur above the mega-ripple and become increasingly burrowed upwards, terminating in a 100% bioturbated sandstone; trace fossils recognised are diplocraterion habichi, skolithos isp. and thalassinoides suevica. pelion formation, locality 3 (fig. 2). 840 scale wave-ripples and swales in the topset deposits suggests that the water depth was above fair-weather wave base in a shoaling wave environment. wave-eroded intraset-boundaries indicate periods with no net deposition. symmetrical mega-ripples were formed by storm waves, and plane-bedded intrasets were formed by suspension fall-out. large-scale inactivity and reactivation surfaces internally in the clinoform beds indicate lateral shifting of the active depositional surface during deposition, or minor sea-level changes. the upper surface of the clinoform beds was erosionally planed off during transgressive marine erosion succeeding deposition of each unit. marine erosion was followed by marked increase in the water depth, reflected by the presence of a strongly burrowed marine surface of erosion, overlain by burrowed calcareous offshore siltstones of facies 5. although the internal sedimentary structures of the clinoform beds closely resemble those observed on active depositional surfaces of modern offshore tidal sand bars, they are not analogous (surlyk & noe-nygaard 1991). a comparable setting may, however, occur in the modern niger delta, where a tabular, 20 km wide sandy, shoreface attached sand unit occurs at 5–10 m water depth in front of a marine, reworked, prograding delta (allen 1965; oomkens 1974). similar high-angle clinoform beds are characteristic of the volgian raukelv formation of southern jameson land and are interfig. 31. a clinoform abandonment zone showing two levels of biological colonisation (1, 2), separated by cross-bedded, medium-grained sandstones. dominant trace fossils are diplocraterion isp. (arrow 1) and thalassinoides isp. (arrow 2). the dark colour of the bioturbated levels is caused by the presence of carbonaceous debris, mud and carbonate cement. pelion formation, locality 3 (fig. 2). fig. 30. a coincident marine surface of erosion (mse) and flooding surface (fs), at the top of a prograding clinoform complex. vertical diplocraterion habichi and monocraterion tentaculatum descend in large numbers from the surface (arrow 1). the uppermost 0.3 m of the clinoform bed show an upwards increase in the degree of carbonate cementation (calcite, dolomite and ankerite). the mse/fs is overlain by offshore siltstones of the fossilbjerget formation (arrow 2), which are calcareous in the lowermost 0.1–0.2 m. pelion formation, locality 4 (fig. 2). preted as marine-reworked shelf-margin wedges formed during sea-level lowstand (surlyk & noe-nygaard 1995). f. offshore association thin offshore deposits that form part of shoreface associations have been described and interpreted in the shoreface section. thick offshore deposits characterise the fossilbjerget formation, and are described and interpreted below (fig. 32). the offshore association consists of sediments deposited from suspension fall-out, at some levels influenced by wave-agitation and currents (facies 5, 6, 10). marine body fossils are abundant throughout, and are commonly concentrated at surfaces capping coarsening-upwards units. offshore sediments immediately overlying such surfaces may contain zoophycos laminatus? and phoebichnus trochoides. the facies associations of the fossilbjerget formation show a proximal–distal, north–south transition from shoreface and sand sheet associations in northern jameson land (facies 7, 9, 10) to fully offshore associations in central and southern jameson land (facies 10; figs 32, 33). in southern jameson land, the offshore deposits consist of structureless and laminated siltstones and mudstones with bands and lenses of fossil-rich carbonate and phosphate-cemented siltstones. an overall fining-upwards tendency has been recorded in these deposits, and deposition took place well below stormwave base. in central jameson land (fig. 2, locality 5), the offshore association is up to 10 m thick, and generally coarsens upwards, with an overall upwards increase in sand content. the association wedges out towards the south, and is generally less than 1 m thick at localities 3 and 4. it may internally contain coarsening-upwards units, grading upwards from structureless or laminated siltstones to fine-grained, structureless, micaceous, carbonaceous and glauconitic sandstones. some units are very rich in fossils, and their upper parts may contain ammonites, embedded at all angles, and brachiopods and bivalves, commonly in life position. proximal offshore units are exposed at locality 9, where a vertical stacking of offshore, shoreface, and shallow marine, nearshore sand sheet associations occur (fig. 33). the association is interpreted as the fully marine, distal offshore part of progradational sandy shoreface units, deposited near or below storm wave-base by suspension fall-out. the thin coarsening-upwards units may reflect minor progradational events. the overall southwards fining and tapering of the offshore deposits of the fossilbjerget formation reflect sediment influx 841 bo un di ng su rf ac es fa ci es as so ci at io ns se di m en ta ry fa ci es tr ac e fo ss ils bo dy fo ss ils c ur re nt or ie nt at io ns pelion and fossilbjerget formations, locality 3, 386–506 m a. sl clay silt f m sand pebbles c f c mse f e 5 10 9 4 4 5 7 8 9 10 12 15 18 19 120 100 80 60 40 20 0 m fossilbjerget fm pelion fm hareelv fm fig. 32. offshore deposits (association f) of the fossilbjerget formation, locality 3. ammonite horizons are indicated; the numbers refer to the faunal horizons listed in full on fig. 37. for legend, see p. 823. to the offshore environment from the north, and distal condensation. biostratigraphical data show that the distal deposits of southern jameson land are strongly condensed at several levels (fig. 6). the degree of condensation decreases gradually towards the north, and the area of condensation expanded northwards through time; the upper part of the fossilbjerget formation consists mainly of condensed, offshore facies throughout the study area (surlyk et al. 1973, 1993; birkelund 1975; heinberg & birkelund 1984; surlyk 1991, 2003, this volume; callomon 1993). sequence stratigraphy a hierarchy of depositional sequences is recognised in the pelion and fossilbjerget formations (surlyk 1990b, 1991; engkilde & surlyk 1993; surlyk et al. 1993; engkilde 1994). a total of 28 sequences are recognised and are expressed by the facies associations. they are stacked into eight low-order composite sequences (p1–8, fig. 6), each containing 3–5 higher order sequences, formed by shoreface progradation. the composite low-order sequences each cover a time span of 1–2 ma, in the range of the third-order composite depositional sequences of mitchum & van wagoner (1991) and are referred to as such in the following discussion. p1–8 are defined by the stacking pattern of the constituent high-order 842 bo un di ng su rf ac es fa ci es as so ci at io ns se di m en ta ry fa ci es tr ac e fo ss ils bo dy fo ss ils c ur re nt or ie nt at io ns pelion and fossilbjerget formations, locality 9, 1002–1060 m a. sl clay silt f m sand pebbles c f c mse mse mse f a a mse a fs a fs fs f f fs f fs f fs f fs f fs f fs f fs fs mse d d a 6 9 7 6 9 7 7 7 6 6 6 6 6 6 6 6 6 6 6 7 1 1 9 7 7 6 10 60 50 40 30 20 10 0 m 17b 17b 16b 16b 16a 15 sb7/tse hst tst sb6/ tse level of maximum flooding fig. 33. offshore (f) and shoreface (a) associations of the fossilbjerget formation. ammonite horizons are indicated. this section is situated in a more proximal position than that shown in fig. 32, and consists of stacked shallow offshore and shoreface associations. from 1–23 m, the section reflects a backstepping trend (transgressive systems tract, tst), and from 23–54 m a forestepping trend (highstand systems tract, hst), as proximal shoreface units overlie distal offshore units. at 54 m, an abrupt transition occurs from shoreface to offshore units. this section illustrates the interdigitation of the pelion formation with the fossilbjerget formation; the pelion formation lowermost in the section (0–5 m) is succeeded by sediments of the fossilbjerget formation (5–23 m), which, in turn, are overlain by sandy pelion formation deposits (23–54 m) referred to the parnas member (surlyk 2003, this volume, fig. 5). offshore mud-rich sediments of the fossilbjerget formation resume above 54 m. locality 9 (fig. 2). the deposits belong to the composite sequence p6, which is bounded by coincident sequence boundaries (sb) and transgressive surfaces of erosion (tse), sb6/tse beneath and sb7/tse above. for legend, see p. 823. sequences and by the nature and extent of their bounding surfaces. the average duration of a high order sequence is about 360,000 years, in the range of the fourth-order depositional sequences of mitchum & van wagoner (1991). they are referred to as sequences or simple sequences. the sequence boundaries of the composite sequences are defined by significant turn-around points in the stacking pattern of the constituent sequences (fig. 6). boundaries of composite sequences are interpreted to represent periods of widespread subaerial exposure in the basin, indirectly reflected by the texture and distribution of overlying lag deposits, areal distribution of underlying shoreface deposits, and in some cases directly by the biostratigraphic duration of the associated hiatuses. the simple sequences form laterally linked sedimentary bodies, bounded by marine surfaces of erosion in shoreface and near-shore successions and correlative conformities in distal, offshore deposits. the marine erosion surfaces, which top most of the sequences in proximal areas are commonly overlain by laterally extensive pebble and sandstone lags (facies 1, 2) formed by transgressive erosion and winnowing of underlying deposits, and the surfaces are thus ravinement surfaces. the lags are generally significantly coarser-grained than the deposits below, implying that relatively large volumes of sediments have been eroded away (engkilde 1994). they represent reworked fluvial, coarse delta-front, or adjacent shoreface deposits reworked during transgression, and their presence suggests that a zone of emergence, sedimentary bypass or erosion was formed after shoreface progradation. the lags thus represent a significant basinwards shift in facies. the lower boundary of a marine lag is accordingly interpreted as a sequence boundary, overprinted by a younger marine ravinement surface. in more distal areas, the sequence boundaries are expressed as submarine erosion surfaces and their correlative conformities. the successions between the sequence boundaries are regarded as small-scale sequences (sensu van wagoner et al. 1990; fig. 34). the top of a lag is interpreted as a marine drowning surface, formed at the time when the water depth became too great for wave reworking. the drowning surfaces do not represent significant hiatuses, but are levels of slow deposition and condensation. they are commonly overlain by laterally widespread mudstones, marls or fine-grained sandstones deposited below storm wave base during the time of maximum flooding. overlying shoreface deposits, forming the bulk of the sequences in proximal areas, are generally 5–15 m thick, and represent southwards progradation over several hundred kilometres without significant changes in thickness and facies. a general thickening of the most basinally positioned parts of the shoreface/tidal inlet units occurs, however, and they may reach up to 35 m in thickness at localities 5, 6, 6a and 7. the basin deepened gradually towards the south, along the basin axis, as indicated by the large-scale facies development. the shoreface/ tidal inlet units are of approximately equal thickness 843 sedimentary bypass and/or erosion following deposition, 4th order sb/tse locally thicker marine sand sheet deposits of the tst transgressive lag, tst offshore mud, tst and mf tst initial basin profile relative sea level through time 1 1 3 3 2 2 44 n s hst fsst sb lst fig. 34. relative position of the systems tracts of the high-order sequences in the jameson land basin. the north–south extent of the progradational systems ranges between 150 and 200 km; the dip of depositional slopes is exaggerated. the relative sea-level stand and the lateral translation of the depositional system are indicated schematically for times 1–4. progradation took place during slow rise or stillstand (time 1), followed by fall (time 1–2), and terminated by slow rise of relative sea level (times 2–3). backstepping took place during the rise in relative sea level (times 3–4). it has not been possible to separate the highstand (hst) and the falling stage (fsst) systems tracts (terminology of plint & nummedal 2000), as there are no significant differences in vertical position. this is due to the low-gradient basin geometry, and to wave-ravinement processes active during transgression. in the field, sandstones of the late lowstand systems tracts (lst) tend to be significantly thicker than those of the hst and fsst. sb, sequence boundary; tse, transgressive surface of erosion; tst, transgressive systems tract; mf, level of maximum flooding. throughout most of northern jameson land, except in the easternmost part where they are thicker. in central jameson land, the signature of the sequences is gradually lost, and the depositional pattern of the composite sequence is dominant. the overall geometry of the sequences is interpreted to reflect progradation during stillstand, fall, and early rise of relative sea level (fig. 34). the bulk of the sandstone part of the sequences is accordingly interpreted to represent falling stage system tracts (terminology of plint & nummedal 2000), as the sandstones are relatively thin over large areas, without any offshore thickening, and as they are capped by complex erosion surfaces. the thicker distal parts of the sequences, composed of sandstone, are interpreted as lowstand systems tracts, deposited during stillstand and early rise in relative sea level. the lowstand systems tracts are attached to the falling stage systems tracts (fig. 34). when the rate of relative sea-level rise became larger than the rate of sediment accumulation, the shoreface system rapidly backstepped towards the basin margins and up to several hundred kilometres northwards up the basin axis. early highstand deltaic deposits are thus to be expected to occur in the marginal areas. composite depositional sequences the eight composite depositional sequences of the pelion–fossilbjerget couplet are described from below (p1–8), including biostratigraphy, boundary relationships, large-scale depositional geometries (systems tracts), and internal architecture. constituent higher order sequences are indicated by a letter suffix, p1a, p1b, etc. all systems tracts of a depositional sequence are not necessary represented at all localities, and successive systems tracts do not necessarily stack vertically (posamentier & vail 1988; posamentier et al. 1988). throughout northern jameson land, the composite sequences are highly asymmetric, volumetrically dominated by the highstand systems tracts (figs 35, 36). this is thought to reflect the low-gradient basin topography, high sediment influx and efficient sediment transport. after filling of the accommodation space in 844 s n 1 2 3 4 5 8 10 11 sb3/tse pelion formation sortehat formation sb3 sb1 mfs tst lst tse lst sb2 sb3 hst hst hst sb2/tse hst sb1 sb1/tse p1 p2 p2 p1 c. 30 m fig. 35. schematic reconstruction of the depositional pattern of the composite sequences p1 and p2, based on sedimentological logs (indicated by vertical lines) and biostratigraphical data (figs 37a, b); the section represents a lateral s–n distance of c. 120 km. the yellow colour indicates shallow marine fineto coarse-grained sandstones, parallel lines represent mainly offshore laminated siltstones and fine-grained sandstones. the sandstone units on the right side of the diagram are shoreface and estuarine units, forming high-order sequences (p1a–d and p2a–e). the estuarine sandstones are interpreted to fill minor incised river valleys and were deposited during p1 late lowstand and early transgressive times. the northwards extent of the sandstones of the lowstand systems tracts (lst) of p1 and p2, on the right side of the diagram, is hypothetical, due to lack of exposures between localities 4 and 5. the coarse grain size of the sandstones in the lst of sequence p2 suggests that the deposits were fed directly by rivers, or linked to a shoreline to the north, and formed a laterally widespread submarine deltaic platform. sb, sequence boundary; tse, transgressive surface of erosion; tst, transgressive systems tract; mfs, maximum flooding surface; hst, highstand systems tract. proximal areas during early highstand, a wide subaerial bypass zone was developed during late highstand, and rivers transported most sediment volumes to the rapidly prograding shorefaces. the shoreface successions were deposited during repeated higher order sea-level falls, and are stacked vertically in a forewards-stepping manner throughout northern jameson land. further offshore in central jameson land, the higher order pattern is less distinct, and very thick coarsening-upwards offshore–shoreface units represent the distal parts of the highstand systems tracts of the composite sequences. the bulk of the transgressive systems tracts of the composite sequences are predicted to exist in the most marginal areas, in milne land to the west and in the western and northern parts of the traill ø – geographical society ø – hold with hope area to the north (fig. 36). sequence p1,upper bajocian (figs 6, 35, 36, 37a, 38a) p1: basal sequence boundary, sb1 in jameson land, the basal sequence boundary (sb1) is of late bajocian age, corresponding to the base of the c. borealis chronozone (figs 6, 37a; surlyk et al. 1973, 1993; surlyk 1990a, b, 1991, 2003, this volume). this sequence boundary is present throughout central east greenland. the ammonite fauna (cranocephalites borealis) belongs to the boreal faunal realm, which cannot be directly correlated with the tethyan faunal realm. the c. borealis chronozone is referred to the upper bajocian (callomon 1961, 1993). it is stressed that correlation to the european bajocian–bathonian stages is uncertain and the recognition of the lower and upper 845 p8 p7 p6 p5 p4 p3 p2 p1 p. athleta e. coronatum k. jason s. calloviense p. koenigi c. nordenskjoeldi c. apertum c. calyx c. variabili a. cranocephaloide a. ishmae a. greenlandicus a. arcticus c. pompeckji c. indistinctus c. borealis chronozones southern jameson land central jameson land northern jameson land traill ø and geographical society ølocalities 1, 2 localities 3, 4 locality 5 localities 6–11 localities 12–14 top fossilbjerget formation, condensed unit lst ? < 80 km3 lst ?, hypothetical lst ?, hypothetical lst ?, hypothetical lst ?, hypothetical lst ?, hypothetical lst 80–110 km3 lst 180–240 km3 hst 1200–1900 km3 hst 1200–1400 km3 tst 1600–1900 km3 tst 1200–1400 km3 tst 900–1000 km3 tst 900–1000 km3 hst 900–1000 km3 hst 900–1000 km3 tst 700–800 km3 tst 400–500 km3 hst 400–500 km3 hst 400–500 km3 tst 400–500 km3 hst 400–500 km3 hst 400–500 km3 tst ??? km3 fig. 36. distribution of pelion formation sandstones in time and space in the jameson land basin. estimates of sandstone volumes (yellow) in the different third order systems tracts are indicated. inferred lowstand deposits are shown (also with estimated volumes, where feasible). the total sandstone volume in the formation is estimated to be in the range of 12,000–15,000 km3. the durations of the highstand and the transgressive systems tracts in a composite sequence are postulated to be of the same order. thick deposits of the transgressive systems tracts are predicted to have been deposited in the traill ø – geographical society ø region. no major variation in the rate of sediment influx is expected to occur at the scale of a single sequence, and the sandstone volumes of the two systems tracts are accordingly estimated to be of the same order. 846 pe lio n an d fo ss ilb je rg et f m s so rt eh at f or m at io n so rt eh at f m sb3 sb3 sb2 p1 sb1 p2 sb1 p1 sb2 p2 a 1 3 4 5 7 8 9 10 11 7 km5 km5 km28 km15 km 4 km 2 km25 km10 km25 km 6a6 northsouth jameson land 9a 2 1b 1a 1a 1b 6 3 4 4 4 5 5 7 8 9a 9 9 8 8 8 8 6 5 5 6 1 5 1 1 1b 1b fossilbjerget fm pelion fm tse sb tse tse sb/tse sb/tse sb/tse sb/tse 10 m 2e 2d 2c 2b 2a 1d 1c 1b 1a ? ? ? ? ? ? 1a, b c. borealis α and β 2 c. indistinctus 3 c. intermissus 4 c. gracilis 5 c. carlsbergensis 6 c. furcatus 7 c. pompeckji 8 c. episcopalis 9 a. arcticus 9a a. delicatus 9b a. micrumbilicatus 10 a. greenlandicus 10a a. freboldi 11 a. harlandi 12a a. ishmae α 12b a. ishmae β and/or a. crassiplicatum 13 a. cranocephaloide and/or k. tychonis 14 k. tychonis and/or k. rosenkrantzi 15 k. peramplus and/or k. vardekloeftensis 16a c. apertum α 16b c. apertum β 17a c. nordenskjoeldi α 17b c. nordenskjoeldi β 18 p. koenigi chronozone 19 s. calloviense 20 k. jason 21 l. keyserlingi planar cross-bedding third order sequence boundary coinciding with a transgressive surface of erosion correlative conformity fourth order sequence boundary transgressive surface of erosion tse sb; sb/tse formation boundary middle to upper shoreface coarseand medium-grained sandstones offshore and lower shoreface fine-grained sandstones and siltstones trough cross-bedding planar lamination, hummocky cross-stratification cross-lamination indistinct lamination structureless ammonite faunal horizons major facies and key surfaces 847 fossilbjerget fm pelion fm hareelv fm h ar ee lv f or m at io n pe lio n an d fo ss ilb je rg et f or m at io ns olympen fm 1 3 4 5 7 8 9 9a 10 11 7 km3 km2 km5 km28 km20 km25 km10 km25 km northsouth jameson land p8 p7 sb7 sb6 p5 sb5 p4 sb4 p3 sb3 p2 p6 ? sb7 sb6 sb5 sb4 sb3 19 19 20 20 15 15 18 14 1312 12 13 14 15 16 19 20 20 18 18 18 19 1919 16 16 15 15 13 13 11 10 11 11 10 9b 9 8 9b 9b 8 8 21 11 10 10 10 9b 9 9 9 6b 6a 5e 5d 5c 5b 5a 4d 4c 4b 4a 3e 3d 3c 3b 3a 6c 6d 6e 12 sb/tsetse sb/tse sb/tse 10m b 5 10 6 6a 7 9 9a 8 1 3 4 11 30 km n a b 12b 12a 12b 17b 16b 16a 12b 9a 10a fig. 37. correlation diagrams for the composite sequences p1–8, and their constituent higher-order sequences in jameson land. the localities and correlation lines are shown on the inset map. predicted extents of the sequences in a northwards direction are shown schematically in fig. 36. correlations are based on ammonites collected by the authors and identified by j.h. callomon (personal communications 1992, 1993, 1994), ammonite collections by t. birkelund and c. heinberg (birkelund 1970, 1971; birkelund & heinberg 1974), data from heinberg & birkelund (1984) and callomon (1959, 1961, 1993), j.h. callomon (personal communications 1994, 1996), and visual correlations in the field. a: composite sequences p1 and p2. the transgressive surface of erosion (tse) that coincides with sequence boundary sb2 in the northern areas is chosen as the datum for the correlation, as this surface is thought to have been essentially flat, or dipping only very slightly in a southwards direction. south of locality 3, this tse passes into a nonerosional, marine flooding surface characterised by the presence of c. gracilis or c. carlsbergensis throughout the region. b: composite sequences p3–8. the transgressive surface of erosion (tse) that coincides with sb4 in the northern areas (locality 5 and northwards) is chosen as datum for the correlation of the northern sections, as it is easily identified in the field and occurs in one faunal horizon (a. micrumbilicatus) throughout the area. in central and south jameson land, the tse that coincides with sb5 at locality 5, is chosen as the datum, as it is a very distinct surface in the field. at locality 5, the surface forms the boundary between the pelion and fossilbjerget formations. south of locality 5, the surface is a non-erosional flooding surface, occurring in one ammonite subzone (a. ishmae β) throughout the area. some of the numbers referring to faunal horizons 9, 12 and 16 do not have a suffix; this indicates that the ammonite in question may be one of two variants, α or β. bajocian and bathonian in east greenland is a best estimate of callomon (1993). the internal zonal correlation within the framework of the boreal middle jurassic ammonite zonation is, however, very precise (surlyk et al. 1973; callomon 1993). unpublished 87sr/86sr values obtained from belemnites of the c. borealis chronozone indicate a late bajocian age, by comparison with the 87sr/86sr isotope curve of jones et al. (1994). in jameson land, the pelion formation overlies the sortehat formation. 87sr/86sr data indicate that the sortehat formation at least in the southern and central part of jameson land includes the aalenian – lower bajocian. dinoflagellate data suggest that the sortehat formation covers the aalenian and possibly part of the bajocian (underhill & partington 1993; koppelhus & hansen 2003, this volume). further north, on traill ø, the pelion formation rests on upper triassic redbeds of the flemming fjord formation (clemmensen 1980), indicating a northwards increase in the duration of the hiatus expressed by sb1 (fig. 6). on northern hold with hope, the pelion formation overlies the lower triassic wordie creek formation (stemmerik et al. 1997; vosgerau et al. in press a). in central and northern jameson land (fig. 2, localities 5, 6, 6a, 7, 9, 11), sb1 is interpreted as a subaerial unconformity, coinciding with a transgressive surface of erosion (tse). it appears as a sharp boundary, separating dark grey, shaly offshore siltstones of the sortehat formation from yellow, burrowed, wave-rippled shoreface sandstones, or laminated light grey, burrowed offshore siltstones and sandstones of the basal pelion formation. mudstone clasts of the sortehat formation locally occur in the lower part of the pelion formation. sb1 is overlain by thin floodplain and estuarine deposits at locality 10, and the sequence boundary here coincides with the base of an incised valley (figs 22, 23, 35, 37a). in southern jameson land, sb1 is represented by a marine regressive surface of erosion, which marks a significant basinwards shift of facies. sb1 here separates dark, silty mudstones of the sortehat formation from shallow marine coarsegrained siltstones and fine-grained sandstones of the pelion formation, which form a lowstand wedge. p1: lowstand systems tract estuarine deposits, up to 40 m thick, are preserved at locality 10, in northern jameson land, and are interpreted to form part of the lowstand systems tract. the deposits fill a valley incised into the sortehat formation during falling and low sea-level stand. the base of the valley forms sb1 (figs 22, 23, 35, 37a). the estuarine deposits may belong to the c. borealis chronozone, but no ammonites were found. the incised valley was formed by fluvial erosion, as indicated by the basal erosional remnants of floodplain deposits. possible correlative fluvial conglomerates of the bristol elv formation occur below marine sandstones of c. borealis chron age on traill ø and probably occupy a major incised valley complex (therkelsen & surlyk in press). in southern jameson land (fig. 2, localities 1–4), southwards thinning basal silty sandstones of p1 form a 20–80 m thick lowstand systems tract of late bajocian age (figs 35, 36, 37a). it consists of heavily burrowed fineto medium-grained sandstones, with bivalves, gastropods, belemnites, ammonites and the trace fossils diplocraterion parallelum, diplocraterion habichi and planolites isp. petrified wood occurs in large quantities. the sandstones are generally poorly cemented. the upper part of the lowstand systems tract consists of sandstones showing well-preserved smalland largescale wave ripples, indicating deposition in a shallow marine, wave-influenced environment, above fairweather wave base, and is interpreted as comprising aggradational and progradational shoreface and shallow marine near-shore deposits. the lower boundary of the lowstand systems tract is an erosion surface, with a relief of up to 0.3 m. the upper boundary is sharp and planar, and is covered by an abundance of belemnites, ammonites (c. borealis α), oysters and bivalves. large numbers of diplocraterion habichi descend from the surface. overlying deposits consist of offshore siltstones of the fossilbjerget formation at localities 1 and 2 and silty sandstones of the pelion formation at localities 3 and 4 belonging to the transgressive systems tract. during the time span of the late lowstand and transgressive systems tracts, the active shoreface rapidly backstepped in a northwards direction. the sandstone volume of the lowstand systems tract is estimated at approximately 3 km3/km width in an e–w section. the complete width is not known, but a conservative estimate is 60–80 km, giving a total sandstone volume in the range of 180–240 km3 (fig. 36). p1:transgressive systems tract the lowstand systems tract is capped by a marine erosion surface throughout southern and central jameson land, which forms the base of the transgressive systems tract. the transgressive systems tract is less than 6 m thick in south jameson land (fig. 2, localities 1, 2), measured from the top of the lowstand deposits to the first occurrence of the c. borealis β faunal horizon (fig. 35) 848 849 milne land traill ø liverpool land geographical society ø hold with hope hold with hope clavering ø 72°n 74°n jameson land a late bajocian milne land traill ø geographical society ø clavering ø 100 km100 km 100 km jameson land b late bathonian milne land traill ø geographical society ø jameson land c middle callovian clavering ø 24°w 20°w basin margin shallow marine and fluvial sandstone and conglomerate shallow marine sandstone offshore silty mudstone offshore mudstone main direction of sediment transport 28°w hold with hope 72°n 74°n 24°w 20°w28°w 24°w 20°w28°w 72°n 74°n fig. 38. palaeogeographic maps showing the large-scale facies distribution in the jameson land basin. a: late bajocian, p1 time. laterally widespread, sandy lowstand deposits characterise southern jameson land during early p1 times. fluvial bypass probably prevailed throughout northern and central jameson land. only a single outlier occurs in liverpool land; it shows sandy deposits of roughly this age, and localities in the easternmost part of the jameson land basin show no indications of sediment influx from the east. modified from surlyk (1977b). the eastwards opening of the rifted seaway in se traill ø is based on carr (1998) and vosgerau et al. (in press b). b: late bathonian, p5 time. throughout the bathonian, southern jameson land was characterised by offshore siltstones whereas central and northern jameson land was dominated by shoreface sandstones. c: middle callovian, p8 time. southern, central and most of northern jameson land was dominated by offshore siltstones. only the northernmost part of jameson land was characterised by shoreface sandstones. and consists of dark grey fossiliferous, phosphatic and calcareous siltstone. at localities 3 and 4, the transgressive systems tract is 30 m thick, contains the c. borealis α fauna throughout, and consists mainly of dark grey, calcareous and phosphatic, laminated offshore siltstones. only minor vertical changes in texture characterise these deposits, suggesting that deposition kept pace with sea-level rise. an overall upwards increase in the density of fossils is seen in the systems tract at localities 3 and 4, suggesting a decrease in sedimentation rate, an increase in biological production or simply better preservation. a general high degree of carbonate cementation characterises the deposits, in very sharp contrast to the underlying lowstand and overlying highstand deposits. in northern jameson land, the transgressive systems tract is represented only by a transgressive surface of erosion, a thin transgressive lag, or in some cases a thin transgressive sand sheet. the c. borealis α fauna is not found north of locality 4. at locality 5, the lower part of the pelion formation is characterised by the c. borealis β fauna. a presumed specimen of c. borealis β (a microconch) was found immediately above sb1 at locality 6, suggesting that the lack of the c. borealis α in the northern part of the study area is primary, reflecting the northwards onlap of the pelion formation. at localities 3 and 4, the transgressive systems tract is capped by a phosphatic, fossiliferous, calcareous siltstone bed, 0.2 m thick. the bed contains large numbers of belemnites, ammonites and bivalves, and is interpreted as a condensed unit, representing the time of maximum flooding (fig. 30). the top surface of the bed is covered with c. borealis β. at localities 6, 6a and 7a, a heavily burrowed, calcareous siltstone bed is situated immediately above the basal sb1/tse, and also represents the time of maximum flooding. p1: highstand systems tract the highstand systems tract of p1 downlaps onto the upper condensed deposits of the transgressive systems tract. it contains c. borealis β at locality 6 in eastern central jameson land, and throughout southern jameson land (fig. 36). the deposits change from proximal, waveand current-influenced shoreface and tidal inlet deposits, c. 120 m thick in the north (fig. 2, localities 5, 6, 6a, 7, 9–11), to a shallow offshore, storm-influenced coarsening-upwards siltstone and sandstone unit, 20–30 m thick, further south (localities 3 and 4). most distally, at locality 1, the highstand deposits consist of a few metres of laminated offshore siltstones (figs 35, 37a). at localities 3 and 4, the highstand systems tract is poorly cemented, and shows a significant upwards decrease in fossil content, in sharp contrast to the underlying deposits of the transgressive systems tract; the highstand systems tract is capped by a distinct marine erosion surface, covered with belemnites, ammonites and the bivalve camptonectes broenlundi (ravn 1911). the surface marks a major basinwards shift in facies, and is a sequence boundary, sb2. diplocraterion habichi descends in large numbers from the surface, and the bivalve modiolus strajeskianus (d‘orbigny 1845) is found in life position in high densities just below the surface, indicating an oxic, shallow marine environment (fürsich 1984). deposits of the highstand systems tract underlie the middle part of the ugleelv member at localities 3 and 4. throughout northern jameson land, the highstand systems tract includes four simple sequences (p1a–d), which form a forwards-stepping sequence set (figs 6, 35, 37a). from locality 5 and northwards, at least as far as locality 11, the upper parts of the sequences show an increase in maximum grain size through p1a–d, reflecting the overall forestepping stacking order. in northern jameson land, the fourth order sequences are capped by coincident sequence boundaries and transgressive surfaces of erosion (sb/tse), which are laterally most extensive and extend furthest offshore in sequence p1d. the maximum progradation point of shoreface units shifted more than 30 km towards the sse during the time span covered by the highstand systems tract of p1 (from near locality 8 in northern central jameson land to south of locality 5 in central jameson land). the volume of the sandstones of the highstand systems tract in jameson land falls in the range of 1100–1400 km3 (deposits may have extended further to the east and west than the present outcrop belt), and the total sandstone volume of the systems tract is estimated at 1600–1900 km3, assuming that sandy highstand deposits continue 120 km northwards into the traill ø – geographical society ø area in an up to 100 m thick, 80 km wide, northwards tapering and thinning wedge (fig. 36). in northern geographical society ø, the stratigraphic interval equivalent to p1 is less than 30 m thick, from the basal onlap surface (sb1) to the first occurrence of cranocephalites indistinctus which coincides with the base of p2 in jameson land. the section consists of a single unit coarsening-upwards from fineto coarsegrained pebbly sandstone (a.g. whitham, personal communication 1996). 850 sequence p2, upper bajocian (figs 6, 35–37a, 39) ammonite evidence indicates that p2 spans the upper bajocian c. indistinctus and c. pompeckji chronozones (callomon 1993). unpublished 87sr/86sr data obtained from belemnites also indicate a late bajocian age, by comparison with the strontium isotope curve of jones et al. (1994). p2: basal sequence boundary, sb2, and lowstand systems tract the basal sequence boundary, sb2, is developed as a distinct surface at localities 3 and 4, and separates the c. borealis and c. indistinctus chronozones. at localities 1 and 2, sb2 is a correlative conformity (figs 35, 37a). in northern jameson land, sb2 coincides with a transgressive surface of erosion (tse), and is commonly overlain by a transgressive lag. at localities 3 and 4, sb2 is overlain by shallow marine fine-grained sandstones, followed by mediumto coarse-grained marine clinoform-bedded sandstones, in all 35–40 m thick. the sandstones are interpreted as a lowstand systems tract, spanning the c. indistinctus and lowermost c. pompeckji chronozones (figs 6, 27, 37a). the deposits form the ugleelv member at localities 3 and 4. a distinct transgressive surface of erosion caps the systems tract, and passes northwards into sb2/tse (fig. 35). the hiatus corresponding to sb2 in northern jameson land is documented at localities 6 and 6a, where p1 contains the c. borealis β fauna in the uppermost part, directly overlain by the c. carlsbergensis fauna of the c. pompeckji chronozone (callomon 1993), occurring in the basal transgressive lag of p2 (figs 27, 37a). sb2 thus spans the c. indistinctus and lowermost c. pompeckji chronozones in this area. the lowstand systems tract thins out rapidly just east of locality 3 and a few kilometres southeast of locality 4. the termination in other directions is not exposed. the sandstone volume in the wedge is estimated at approximately 1.4 km3/km width, in an e–w section. the complete width is not known, but is estimated to be 60–80 km, giving a total sandstone volume in the range of 84–112 km3 (fig. 36). p2: transgressive systems tract the lowstand systems tract at localities 3 and 4, and sb2 in northern jameson land, are capped by a transgressive surface of erosion (tse), which passes into a simple flooding surface in southern jameson land. the tse may be present on southern traill ø, as a marine erosional surface overlain by the ammonite c. gracilis (a.g. whitham, personal communication 1995). the tse is overlain by large numbers of c. carlsbergensis at localities 6, 6a, 11–14. the occurrence of the c. gracilis fauna on the surface at localities 3 and 4, and of the slightly younger c. carlsbergensis fauna further north (except possibly in the traill ø region), may reflect diachronous onset of sedimentation on this surface. the surface was formed by erosional retreat of the shoreface depositional system over more than 200 km, from south of locality 5 to somewhere in the traill ø – geographical society ø area, before the onset of renewed shoreline progradation, which ended at least 50 km north of the toe of the underlying lowstand systems tract of p2 (figs 35, 37a). the tse is overlain by a transgressive lag followed by thin offshore siltstones from locality 5 and northwards. it is directly overlain by offshore calcareous siltstones at localities 3 and 4. the total volume of the sandstone part of the transgressive systems tract is broadly estimated to equal the sandstone volume of the preceding highstand systems tract, in the range of 1600–1900 km3 (fig. 36). p2: highstand systems tract the highstand systems tract of p2 is wedge-shaped and belongs to the upper part of the c. pompeckji chronozone. it is up to 90 m thick in northern and central jameson land, and thins to less than 5 m in southern jameson land. it consists of four to five stacked simple sequences (figs 35, 39), which form an aggradational to forwards-stepping sequence set (fig. 35). in northern jameson land, the upper four sequences are generally amalgamated, with very minor or no offshore or lower shoreface deposits, suggesting a low rate of creation of new accommodation space. all sequences prograded to a position south of locality 5. the c. episcopalis fauna is found in the middle to upper part of the systems tract throughout northern jameson land (fig. 37a). the sandstone volume of the systems tract is in the range of 700–900 km3 in jameson land. the total sandstone volume, including deposits in the northernmost part of the basin is estimated at 1200–1400 km3, using the same assumption as for p1 (fig. 36). sequence p3, lower bathonian (figs 6, 36, 37b, 39) ammonite evidence suggests that p3 is of early bathonian age. unpublished 87sr/86sr values from belemnites suggest an early–middle bathonian age, by comparison with the strontium isotope curve of jones et al. 851 (1994). p3 contains the arctocephalites arcticus fauna of the a. arcticus chronozone throughout the study area. in south jameson land, the species a. delicatus of this zone is also found. sb3: basal sequence boundary and lowstand systems tract the boundary between the c. pompeckji and the a. arcticus chronozones coincides with a third order sequence boundary, sb3 (fig. 37b). sb3 coalesces with a tse from locality 5 and northwards. at locality 8, sb3 is interpreted locally to form the base of a minor incised valley, filled with estuarine deposits. sb3 passes into a correlative conformity within offshore deposits of the fossilbjerget formation south of locality 5, and marks a major shift in the lateral distribution of accommodation space at a composite sequence level. the depocentre of p3 is located in northern jameson land, and the thickness of p3 decreases both southwards and northwards of the depocentre (fig. 37b). this contrasts with the underlying p2, which shows a tabular geometry throughout the northern localities southwards to locality 5, south of which it thins rapidly. a single occurrence of estuarine deposits (locality 8), positioned immediately above sb3, is interpreted to form part of the lowstand systems tract. lowstand deposits may have existed between localities 4 and 5, but if so, they have been removed by recent erosion. p3: transgressive systems tract the transgressive systems tract is represented by a tse overlain by lag deposits, which may be followed by thin, calcareous siltstones interpreted to have been deposited during maximum flooding. the upper part of the estuarine association of locality 8 may also belong to the transgressive systems tract. the total volume of the sandstone part of the transgressive systems tract is estimated at 1200–1400 km3, using the same assumptions as for p2 (fig. 36). p3: highstand systems tract the highstand systems tract volumetrically dominates p3 throughout central and northern jameson land. it is 42 m thick at locality 11, thickens to 50–60 m at locality 8 (the thickness is approximate, due to limited stratigraphic control), and thins to less than 15 m at locality 5. sequence p3 is not clearly distinguishable in southern jameson land, but the stratigraphic interval corresponding to the sequence covers less than 8 m at localities 3 and 4. at localities 5 and 8, only the uppermost part of p3 is represented by shoreface deposits, whereas the lower part consists of offshore siltstones and fine-grained sandstones, reflecting the forwards-stepping nature of the sequence set (fig. 37b). during highstand time, shoreface units stepped forwards over more than 30 km. the highstand systems tract is capped by a low order sequence boundary, sb4, which is developed as an unconformity and coincides with a transgressive surface of erosion from central jameson land and further north (fig. 37b). the sandstone volume of the systems tract 852 fig. 39. stacked shoreface and shallow, near-shore sand sheet units of p2 and p3 (fig. 6). sb2/tse and sb3/tse are sequence boundaries, each coinciding with a transgressive surface of erosion, capping sequences p1 and p2, respectively, throughout northern jameson land. tidal inlet channel sandstones (tc) of sequence p1 are seen in the lower part of the section. sequence p2 is 90 m thick and consists of four shoreface units; sequence p3 is dominated by sand sheet deposits. the dark bands are horizons cemented with ankerite, dolomite, calcite and siderite. pelion formation, locality 7 (fig. 2). 853 in jameson land is in the range of 400–500 km3, and the total volume of the sandstone part is estimated at 900–1000 km3, using the same assumptions as for p1 (fig. 36). a carbonate-cemented, fossiliferous, fully marine sandstone, c. 30 m thick, covers the stratigraphic level of p2 and p3 on northern geographical society ø. the unit is overlain by deposits containing a. micrumbilicatus and is interpreted as a tidal channel fill overlain by transgressive sand sheet deposits (a.g. whitham, personal communication 1996). sequence p4, lower–middle bathonian (figs 6, 36, 37b, 38b) ammonite evidence indicates an early–middle bathonian age for sequence p4. unpublished 87sr/86sr values from belemnites also suggest an early to middle bathonian age, by comparison with the strontium isotope curve of jones et al. (1994). p4 contains the a. micrumbilicatus and a. greenlandicus faunas of the a. greenlandicus chronozone in the lower part, and a. harlandi and the a. ishmae α faunas of the a. ishmae chronozone in the upper part of the sequence. the a. ishmae fauna is also found on traill ø. sb4: basal sequence boundary the upper boundary of p3 is formed by a low order sequence boundary, sb4, which coincides with a tse from locality 5 and northwards, except at a position between localities 10 and 11, where it locally forms the base of an incised channel fill, up to 30 m thick, capped by the tse (fig. 37b). south of locality 5, sb4 is developed as a correlative conformity, and the transgressive surface of erosion passes into a simple marine flooding surface. it is suggested that the hiatus associated with sb4 covers the uppermost a. arcticus chronozone in northern jameson land, separating the a. arcticus fauna of the a. arcticus chronozone and the a. micrumbilicatus fauna of the a. greenlandicus chronozone; the a. delicatus fauna of the a. arcticus chronozone is absent. a. delicatus is only found in southern jameson land, where it occurs in the basal part of p4. at locality 5, an uncertain occurrence of this species is indicated, but this might instead be a specimen of a. arcticus (callomon 1993). in any event, a considerable faunal hiatus is thought to separate the a. arcticus and a. delicatus faunas of the a. arcticus chronozone in east greenland. the missing interval may be represented by the a. spathi, a. porcupinensis and a. arcticus faunas in part of the a. spathi and a. porcupinensis chronozones known from the yukon area, canada (callomon 1993). lowstand systems tract deposits have not been identified in p4. deposits of this systems tract may have been deposited in the area between localities 4 and 5 but, if so, have been removed by recent erosion. p4: transgressive systems tract a significant transgression is interpreted to have occurred after sb4 time, leaving a transgressive lag throughout northern jameson land, and possibly further north. this lag is overlain by a thin sheet of offshore siltstones, locally rich in bivalves, belemnites and ammonites, interpreted as a condensed unit and representing the level of maximum flooding. the transgressive lag and the condensed offshore deposits contain the a. micrumbilicatus fauna of the lower a. greenlandicus chronozone throughout jameson land, whereas the a. delicatus fauna is restricted to southern jameson land. the lower faunas of the a. greenlandicus chronozone are found throughout jameson land. the volume of the sandstone part of the transgressive systems tract is estimated to be in the order of 900–1000 km3, including sandy deposits in the traill ø – geographical society ø area, using previous assumptions (fig. 36). p4: highstand systems tract the highstand systems tract of p4 contains ammonites of the middle to upper a. greenlandicus and a. ishmae zones (a. greenlandicus, a. freboldi, a. harlandi and a. ishmae α faunas). sequence p4 is c. 60 m thick throughout northern and central jameson land, and is volumetrically dominated by a forwards-stepping set of simple sequences, p4a–d (fig. 37b). the thickness of the section which covers the stratigraphic interval of p2–4 on geographical society ø is only 30 m (a.g. whitham, personal communication 1995), compared to 250 m in north jameson land. this large difference is interpreted to be the result of a high degree of sedimentary bypass and erosion in the northern, proximal areas during times of falling sea level and lowstand. the sandstone volume of the systems tract in jameson land is in the order of 700–800 km3; the total sandstone volume is estimated at 900–1000 km3 (fig. 36). sequence p5, middle–upper bathonian (figs 6, 36, 37b) sequence p5 is up to 75 m thick, and consists of an aggradational sequence set (p5a–e). the maximum southwards progradation distance of p5 is c. 25 km less than that of the underlying p4, marking the onset of the large-scale backstepping pattern of the composite sequences from p5 to p8 (figs 6, 37b; surlyk 1990b, 1991). p5 spans the upper a. ishmae, a. cranocephaloide, c. variabile and c. calyx chronozones indicating a middle to late bathonian age-span for the sequence (figs 6, 37b). unpublished 87sr/86sr values from belemnites also indicate a middle to late bathonian age, by comparison with the strontium isotope curve of jones et al. (1994). the a. ishmae β fauna of the a. ishmae chronozone is one of the most common and widespread faunas in the arctic. it occurs throughout jameson land, and has also been found on traill ø (donovan 1953). at this level, the genus oxycerites occurs; it is usually restricted to lower latitudes in the tethyan realm, and indicates a late bathonian age (birkelund et al. 1971). the presence of oxycerites suggests that the transgression that ended deposition of p4 was linked to a eustatic rise in sea level, permitting migration of ammonites between the boreal and tethyan realms. sb5: basal sequence boundary sb5 is positioned in the a. ishmae chronozone, separating the a. ishmae α and β faunas. sb5 is developed as an sb/tse in central jameson land (fig. 2, locality 5) and further north. in central jameson land, sb5 tops the uppermost shoreface deposits of the pelion formation, and is overlain by offshore siltstones of the fossilbjerget formation. in northern jameson land, sb5 separates condensed sequences of p4 from the less condensed sequences of p5. lowstand systems tract deposits have not been identified in p5, but may have existed in the area between localities 5 and 8, and if so, have been removed by recent erosion. p5: transgressive systems tract the transgressive systems tract deposits are generally thin. in southern jameson land, the tract is represented by a marine flooding surface overlain by thin offshore siltstones, interpreted as a condensed unit. from locality 5 and northwards, it is represented by a transgressive surface of erosion, locally overlain by a conglomerate lag or sandstone sheet, up to 4 m thick, with abundant a. ishmae β. these deposits are overlain by thin, laminated offshore siltstones, interpreted to represent the level of maximum flooding. the total volume of the sandstone part of the transgressive systems tract is estimated at 900–1000 km3 (fig. 36). p5: highstand systems tract the highstand systems tract of p5 contains arcticoceras crassiplicatum of the a. ishmae chronozone, and a. cranocephaloide, kepplerites tychonis, k. rosenkrantzi, k. peramplus, k. vardekloeftensis and k. svalbardensis of the a. cranocephaloide, c. variabile and c. calyx chronozones. in northern jameson land, the highstand systems tract of p5 is dominated by progradational shoreface units. the shorelines prograded southwards, at least as far south as locality 8 (fig. 37b). at this locality, p5 consists of more than three coarsening-upwards offshore units of the fossilbjerget formation. at locality 5, the upper part of p5 contains k. peramplus and k. svalbardensis of the c. calyx chronozone. in southern jameson land, the c. calyx chronozone marks the onset of a significant increase in the degree of condensation of the offshore deposits of the fossilbjerget formation (surlyk et al. 1973; callomon 1993). the sandstone volume of the systems tract in jameson land is in the order of 400–500 km3, and the total sandstone volume is estimated at 700–800 km3 (fig. 36). sandy, shallow marine deposits of the transgressive systems tract exist in the traill ø – geographical society ø area, corresponding to the level of p5, although a precise biostratigraphic correlation has not yet been made. sequence p6, upper bathonian – lower callovian (figs 6, 36, 37b) the composite sequence p6 spans the c. apertum and c. nordenskjoeldi chronozones. the duration of p6 is about 1.5 ma, according to the time-scale of gradstein et al. (1994; fig. 6). p6 is backstepping in the lower part and forestepping in the upper part. it has only been studied in detail at locality 9a. sb6: basal sequence boundary sb6 coincides with a transgressive surface of erosion throughout northern jameson land. in this area, the hiatus at sb6 corresponds to the k. vardekloeftensis faunal horizon of the c. calyx chronozone, and this fauna 854 is restricted to southern jameson land (fig. 37b; callomon 1993). the c. apertum α fauna of the c. apertum chronozone is found immediately above sb6. at this stratigraphic level, biostratigraphic correlation with european ammonite faunas is again possible, and the c. apertum chronozone correlates with the c. discus and part of the m. herveyi chronozones (callomon 1993). sb6 is situated close to the base of the callovian in east greenland, and forms a very prominent correlation surface throughout central and northern jameson land. a major northwards retreat of the depositional systems is recorded across the sb6/tse (fig. 6). lowstand systems tract deposits have not been identified in sequence p6; if deposited, they are expected to occur south of locality 8. in southern jameson land, p6 is represented by offshore siltstones, which form part of a condensed unit that spans most of the lower callovian (figs 6, 37b; surlyk et al. 1973; birkelund 1975; surlyk 1990b, 1991; callomon 1993). the distal condensation reflects the overall large-scale backstepping of the third order sequences during this time interval. p6: transgressive systems tract at locality 9a, the transgressive systems tract consists of a basal lag, overlain by offshore siltstones and very fine-grained sandstones, 22 m thick (level 1–23 m in fig. 33). the deposits include a transgressive sand sheet, 5 m thick, overlain by a shoreface unit, 5 m thick, which is followed by an overall backstepping set of shallow offshore, storm-influenced coarsening-upwards units, each 1–1.5 m thick, representing the distal edges of shoreface units. at localities 8 and 11, similar deposits are found at the same stratigraphic level, and are interpreted as belonging to the transgressive systems tract. the total volume of the sandstone part of the transgressive systems tract is estimated at 700–800 km3, including sandy deposits in the traill ø – geographical society ø area (fig. 36). p6: highstand systems tract at locality 9a, a change in stacking pattern, and the reappearance of shallow water sandstones in the section, marks the transition from the transgressive to the highstand systems tract. at locality 9a, the highstand systems tract of p6 is 31 m thick and consists of coarseningupwards offshore and shoreface units, separated by non-erosional or erosional marine flooding surfaces (23–54 m in fig. 33). the units are stacked in a forwardsstepping pattern. the uppermost three units show intense burrowing by diplocraterion habichi, monocraterion tentaculatum, ophiomorpha nodosa and skolithos isp., suggesting very shallow water depths. a shallow water depth is also indicated by the occurrence of both small and large wave-ripples as well as swaley cross-stratification. the package comprising the three coarsening-upwards units forms the ‘minor regressive cycle’ of heinberg & birkelund (1984); this unit forms the parnas member of the pelion formation (surlyk 2003, this volume, fig. 5). the deposits are interpreted as near-shore sand sheet and shoreface units. in northernmost jameson land, p6 consists of shoreface deposits only. the sandstone volume of the systems tract in jameson land is estimated at 200–300 km3, and the total sandstone volume is estimated at 400–500 km3 (fig. 36). sequences p7 and p8, lower–upper callovian (figs 6, 36, 37b, 38c) sequence p7 covers the k. koenigi chronozone, with a duration of about 1 ma, and sequence p8 covers the s. calloviense, k. jason, e. coronatum and lower p. athleta chronozones, with a total duration of about 2.5 ma, according to the time-scale of gradstein et al. (1994). sb7: basal sequence boundary sb7 occurs near the boundary between the c. nordenskjoeldi and k. koenigi chronozones. it marks the maximum progradation of highstand deposits of p6, which form the parnas member. in northern jameson land, sb7 is developed as an sb/tse, whereas it is represented by a correlative conformity in central and southern jameson land. the deposits of p7 and p8, lower–upper callovian the deposits of p7 and p8 are not subdivided into systems tracts, due to a lack of distinct depositional trends. the sequences form the uppermost part of the pelion–fossilbjerget couplet. throughout south and central jameson land, the sequences are represented by offshore siltstones of the fossilbjerget formation, grading into shoreface sandstones of the pelion formation in northern jameson land (fig. 37b). a minor progradational unit, 3–4 m thick, is dated to the upper k. koenigi – lower s. calloviense chronozones. it contains the k. galilaeii and s. calloviense 855 faunas (birkelund et al. 1971). in this unit, two sandstone sheets, together up to 3 m thick, separated by bioturbated siltstones, were deposited in the otherwise siltstone-dominated offshore environment (fig. 37b, localities 9a, 11). the sandstone beds show hummocky cross-stratification, and ophiomorpha nodosa burrows, which also occur in the interbedded siltstones. a rich fauna of ammonites, belemnites, bivalves, gastropods and brachiopods occurs in the beds. the trace fossils suggest a nearshore position, and the sandstone beds are interpreted as nearshore storm deposits, laid down above storm wave base. the upper bed is topped by a marine flooding surface, overlain by parallel, even laminated, offshore mudstones. the sandstones are interpreted as the basinwards limit of a forwards-stepping depositional unit, possibly forming the distal edge of lowstand deposits of p8. the top of the sandstones mark the lower–middle callovian boundary in the area. throughout jameson land, p8 forms an important condensed level, which caps the pelion and fossilbjerget formations in the area, and which is downlapped by the olympen formation (figs 6, 37b). in northern jameson land, shallow marine sandstones of the pelion formation correspond to the stratigraphic interval of p7 and p8 (surlyk et al. 1973; heinberg & birkelund 1984). condensation in south jameson land was initiated in p5 time, and in the upper part includes the middle callovian offshore goniomyakløft member of the fossilbjerget formation (surlyk 2003, this volume, fig. 5), the ‘wood beds’ of callomon (1993) containing kosmoceras cf. or aff. jason of the k. jason chronozone (callomon 1961, 1993; surlyk & birkelund 1972; surlyk et al. 1973, 1993; surlyk 1990b, 1991). in central jameson land, longaeviceras keyserlingi of the p. athleta chronozone occurs c. 20 m below the fossilbjerget formation – olympen formation boundary (larsen & surlyk 2003, this volume). in northern jameson land, a single specimen of kosmoceras (zugokosmokeras) cf. phaeinum (buchman) or kosmoceras (zugokosmokeras) cf. proniae (teisseyre) of the p. athleta chronozone was found 2.8 m below the boundary. a fragment of kosmoceras (zugokosmokeras) cf. proniae was reported from 15 m above the boundary in north-west jameson land (birkelund et al. 1971). the goniomyakløft member in southern jameson land thus represents the top of the pelion–fossilbjerget couplet and possibly overlaps with the basal part of the olympen formation to the north. the faunas of the p. athleta chronozone represent a mixture of a true boreal fauna, with longaeviceras, and a sub-boreal kosmoceras fauna (callomon 1993). this faunal mixing strongly suggests that the long term transgression, recorded from p5 to p8, was eustatic in origin. in southern jameson land, the goniomyakløft member is overlain by thin shales of the hades member (olympen formation) or by the shales and massive sandstones of the upper oxfordian – lower volgian hareelv formation (fig. 6; surlyk & birkelund 1972; surlyk et al. 1973; surlyk 1987, 1991, 2003, this volume; callomon 1993; surlyk & noe-nygaard 2001b). regional implications the couplet comprising the pelion and fossilbjerget formations forms the lower half of a large-scale middle jurassic regressive–transgressive–regressive succession, and is approximately time-equivalent with the brent group of the northern north sea and with the krossfjord and fensfjord formations and to some extent the garn formation on the norwegian shelf. the depositional pattern of the brent group has been related to the uplift and subsequent deflation of a major volcanic dome (sellwood & hallam 1974; whiteman et al. 1975; hallam & sellwood 1976; ziegler 1988, 1990; underhill & partington 1993, 1994). the presence of a similar, roughly contemporaneous regional early jurassic uplift and subsequent subsidence in east greenland was suggested by surlyk (1977a, 1978) and surlyk et al. (1993) on stratigraphical grounds and corroborated by fission track thermochronology by johnson & gallagher (2000). the main east greenland uplift was centred in northern east greenland, and the eastern part of the uplift area is today probably located on the norwegian shelf off lofoten. uplift of the east greenland area probably commenced in the early jurassic, culminating in toarcian–aalenian times. the regional uplift is probably partly equivalent to the uplifted intra-rift area between greenland and norway of doré (1991; fig. 5). rifting began in the late bajocian, increased during the bathonian–kimmeridgian and culminated in kimmeridgian – mid-volgian times (surlyk et al. 1981). sequence p1 and the lower part of p2 of the pelion and fossilbjerget formations show the development of shallow marine sandstones of late bajocian – early bathonian age in a basinal position, which are similar to time-equivalent deposits in the northern north sea and the norwegian shelf described, for example, by gjelberg et al. (1987), helland-hansen et al. (1992) and mitchener et al. (1992). major flooding events occurred both in the north sea area and in jameson land near the bajocian–bathonian boundary, indicating the onset of an inter-regional transgression (surlyk et al. 1973, 1981, 856 1993; birkelund 1975; surlyk 1990a, b, 1991, 2003, this volume; helland-hansen et al. 1992; mitchener et al. 1992; steel 1993). the p3 and p4 sequences show a largescale progradational stacking pattern. sb5 at the base of p5 coincides with the lithostratigraphic top of the pelion formation in central jameson land, and marks a major shift from shoreface to offshore depositional environments. sb6 at the top of p5 roughly coincides with the bathonian–callovian boundary in jameson land, and forms the top of the pelion formation throughout northern jameson land, marking a major landwards shift in depositional environments. the p5–8 sequences show a large-scale backstepping pattern. the distance of overall large-scale shoreline backstepping from p1 to p5 is more than 100 km. backstepping through p6–8 is about 50 km. in northern central jameson land, only the upper part of p8 forms part of the condensed succession, which in south jameson land covers most of the callovian, and marks the culmination of backstepping (surlyk 1977b, 1990b, 1991, 2003, this volume; surlyk et al. 1993; alsen & surlyk in press). throughout central east greenland, the condensed unit is characterised by a mixture of true boreal and sub-boreal ammonite faunas (callomon 1993), supporting the sequence stratigraphic interpretation of the inter-regional character of the large-scale transgression. the transgression, which was initiated in the late bathonian (p5), and continued into the early oxfordian, was partly of eustatic origin according to hallam (1988). domal subsidence may also have caused a significant relative sea-level rise through this time span in the areas influenced by the domes, similar to the situation in the north sea (for further discussion, see surlyk 1990b, surlyk et al. 1993 and underhill & partington 1994). summary in middle jurassic times, a north–south elongate shallow marine embayment was formed in central east greenland with a low gradient, without a marked shelfslope break. throughout the east greenland basin, the base of the pelion formation marks the onset of a new depositional cycle characterised by a major re-arrangement of drainage patterns, extensive transgressive onlap and the influx of large amounts of quartzose sand, interpreted as reflecting the onset of the important middle– late jurassic rift phase which culminated in kimmeridgian–volgian times (surlyk 1978, 1990a, b, 2003, this volume; surlyk & clemmensen 1983; surlyk & noenygaard 2000). the low axial gradient of the basinfloor resulted in the development of laterally extensive simple sequences, formed by predominantly axial shoreline progradation. the distribution and architecture of the sandstone units were controlled by changes in relative sea level. a total of 28 simple sequences are recognised. detailed biostratigraphic correlations at the level of ammonite zones and faunal horizons show that the average duration of a simple sequence is about 360,000 years. they are the smallest-scale units which can be correlated over large distances and they are bounded by laterally extensive key surfaces. of special importance for the sequence stratigraphic interpretation is the nature of transgressive marine lags, which form the basal deposits of most sequences. the presence of laterally widespread pebble lags, containing intraand extra-formational clasts, mainly of larger sizes than those present in the underlying deposits, is interpreted to indicate that a modified sequence boundary coincides with the base of the lag deposits (surlyk et al. 1993, 1995). the horizontal distribution of transgressive lag conglomerates directly reflects the original distribution of the former presence of foreshore, beach, and fluvial deposits. the very extensive lateral distribution of the shallow marine sandstones of the pelion formation, over more than 20,000 km2, and the distribution of coarse pebbly sandstones and conglomerates over more than 10,000 km2 indicate that major rivers continuously delivered coarse-grained sediments to the southwards prograding shorelines. the rivers entered the basin at its northern and western margin at relay zones between en echelon right-stepping border faults, as testified by palaeocurrent data and the large-scale textural and facies distribution (surlyk et al. 1973, 1981; surlyk 1977b, 1990a; surlyk & clemmensen 1983). the proportion of conglomerates and coarse-grained sandstones decreases southwards, indicating that the transport capacity of the rivers decreased as the floodplain became wider during progradation. only one localised occurrence of terrestrial deposits is recognised in jameson land. this is most likely a consequence of transgressive marine erosion of floodplain, lagoonal and fluvial sediments of the underlying sequences. the sequences generally consist of a thin transgressive systems tract, dominated by a sandstone or pebble lag overlain by fossiliferous, calcareous siltstones or mudstones representing a condensed unit, followed by sandy highstand, forced regressive and lowstand deposits (fig. 34). detached lowstand shoreface deposits were probably not developed during the high-order cycles because of a continuous, high sediment influx 857 to the basin and a high sediment distribution rate. in proximal areas, the sequences are terminated upwards by coincident sequence boundaries and transgressive surfaces of erosion. in distal offshore areas, the correlative conformities are difficult to pinpoint. definition of low-order composite sequences is based on the stacking pattern of the high-order sequences, the nature and extent of marine transgressive lags, and a detailed biostratigraphical ammonite zonation based on faunal horizons. eight composite sequences are recognised (p1–8), each with a duration of 1–2.5 ma (fig. 6). in proximal realms, the sequences are separated by subaerial erosion surfaces representing sequence boundaries, coincident with marine transgressive surfaces of erosion. more distally, the sequence boundaries are represented by surfaces marking significant basinwards shift in facies (fig. 35). low-order sequence boundaries mark major changes in the overall stacking pattern of the high-order sequences. the composite sequences are subdivided into systems tracts, which in the study area are highly asymmetric, volumetrically dominated by highstand deposits (fig. 35). these consist of aggradational to forwardsstepping sequence sets, generally made up of thin offshore siltstones overlain by progradational shoreface and shallow marine, near-shore sandstones and conglomerates, deposited in response to high-order sea-level fall (forced regressions) and lowstand. the sandy deposits of the low-order highstand systems tract downlap onto extensive condensed offshore deposits, deposited during maximum flooding. two examples of low-order lowstand sandstones are documented, belonging to p1 and p2, while a third, belonging to p8, is less well documented (figs 35, 36). in the northern part of the study area, the transgressive systems tracts are mostly represented by a transgressive surface of erosion, commonly overlain by a thin coarse-grained sandstone or pebble lag or a thin sand sheet. a thicker transgressive systems tract has only been identified in p1, at localities 3 and 4, and in p6 at locality 9 in northern central jameson land. in each sequence, the transgressive systems tract consists of backstepping offshore to shoreface depositional units, reflecting fluctuations in relative sea level on a higher order scale (fig. 33). volumetrically large portions of p7 and p8 may also form part of transgressive systems tracts throughout north–central jameson land, but this has not been quantified due to lack of clear depositional trends. distally, in southern jameson land, the transgressive systems tracts are represented by simple marine flooding surfaces, overlain by condensed offshore deposits. the volumetric distribution of the sandy parts of the low-order systems tracts are quantified based on values obtained from the study area and unpublished data from the traill ø – geographical society ø area. acknowledgements this study is based on field work in jameson land, east greenland during the summers of 1991 to 1995, generously supported by norsk hydro a/s, research centre, bergen. john gjelberg is especially thanked for interest, support and discussion. m.e. gratefully acknowledges the carlsberg foundation (ans. 94-0204/20) for support during the preparation of this paper. the danish natural science research council is acknowledged for support during the last phase of field work. john h. callomon is thanked for determinations of the ammonites collected by the authors, for providing much new material, and for discussion of the biostratigraphy. rené madsen is thanked for drafting, ole bang berthelsen for photography and m. vesterager for word processing. snorre olaussen, john gjelberg and jon r. ineson contributed with much appreciated, highly constructive reviews. references allen, g.p. 1991: sedimentary processes and facies in the gironde estuary: a recent model of macrotidal estuarine systems. in: smith, g.d. et al. 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(eds): the jurassic of north-east greenland. geological survey of denmark and greenland bulletin. whiteman, a.j., rees, g., naylor, d. & pegrum, r.m. 1975: north sea troughs and plate tectonics. norges geologiske undersøkelse 316, 137–161. yang, c.s. & nio, s.d. 1989: an ebb-tide delta depositional model – a comparison between the modern eastern scheldt tidal basin (southwest netherlands) and the lower eocene roda sandstone in the southern pyrenees (spain). sedimentary geology 64, 175–196. ziegler, p.a. 1988: evolution of the arctic – north atlantic and the western tethys. american association of petroleum geologists memoir 43, 198 pp. ziegler, p.a. 1990: geological atlas of western and central europe, 2nd edition, 239 pp. amsterdam: elsevier for shell internationale petroleum maatschappij. manuscript received 6 august 1996; revision accepted 5 may 1998. geological survey of denmark and greenland bulletin 1, 145-216 denmark, southern sweden and the netherlands previous page: jurassic sediments exposed in a clay-pit at bagå on bornholm, denmark – see michelsen et al. (2003, this volume). photo: peter k. warna-moors. 147 jurassic lithostratigraphy and stratigraphic development onshore and offshore denmark olaf michelsen, lars h. nielsen, peter n. johannessen, jan andsbjerg and finn surlyk a complete updated and revised lithostratigraphic scheme for the jurassic succession of the onshore and offshore danish areas is presented together with an overview of the geological evolution. the lithostratigraphies of bornholm, the danish basin and the danish central graben are described in ascending order, and a number of new units are defined. on bornholm, the lower–middle jurassic coal-bearing clays and sands that overlie the lower pliensbachian hasle formation are referred to the new sorthat formation (lower jurassic) and the revised bagå formation (middle jurassic). in the southern danish central graben, the middle jurassic succession formerly referred to the lower graben sand formation is now included in the revised bryne formation. the lulu formation is erected to include the uppermost part of the middle jurassic succession, previously referred to the bryne formation in the northern danish central graben. the upper jurassic heno formation is subdivided into two new members, the gert member (lower) and the ravn member (upper). the organic-rich part of the upper farsund formation, the former informal ‘hot unit’, is established formally as the bo member. dominantly shallow marine and paralic deposition in the late triassic was succeeded by widespread deposition of offshore marine clays in the early jurassic. on bornholm, coastal and paralic sedimentation prevailed. during maximum transgression in the early toarcian, sedimentation of organic-rich offshore clays took place in the danish area. this depositional phase was terminated by a regional erosional event in early middle jurassic time, caused by uplift of the central north sea area, including the ringkøbing–fyn high. in the sorgenfrei–tornquist zone to the east, where slow subsidence continued, marine sandy sediments were deposited in response to the uplift. uplift of the central north sea area was followed by fault-controlled subsidence accompanied by fluvial and floodplain deposition during middle jurassic time. on bornholm, deposition of lacustrine muds, fluvial sands and peats dominated. the late middle jurassic saw a gradual shift to shallow marine deposition in the danish central graben, the danish basin and skåne, southern sweden. during the late jurassic, open marine shelf conditions prevailed with deposition of clay-dominated sediments while shallow marine sands were deposited on platform areas. the central graben received sand by means of sediment gravity flows. the clay sediments in the central graben became increasingly rich in organic matter at the jurassic–cretaceous transition, whilst shallow marine coarse-grained deposits prograded basinwards in the sorgenfrei– tornquist zone. keywords: denmark, danish central graben, danish basin, sorgenfrei–tornquist zone, bornholm, jurassic, lithostratigraphy, basin development o.m., geological institute, university of aarhus, c.f. møllers allé, dk-8000 århus c, denmark. l.h.n., p.n.j. & j.a., geological survey of denmark and greenland, geocenter copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: lhn@geus.dk f.s., geological institute, university of copenhagen, geocenter copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. geological survey of denmark and greenland bulletin 1, 147–216 (2003) © geus, 2003 148 jurassic deposits extend throughout most of the danish area, from the island of bornholm in the baltic sea in the east to the central part of the north sea in the west (fig. 1). they can be studied at outcrop on bornholm and skåne (southern sweden), whereas they are only known from subsurface data over the remainder of the area. the jurassic of bornholm has been studied since the 19th century; the early literature includes forchhammer (1837), jespersen (1866, 1869), bartholin (1892, 1894), grönwall (1899), hjort (1899), møller (1902, 1903), malling & grönwall (1909), malling (1911, 1914, 1920) and höhne (1933). the stratigraphy and structural evolution was studied by gry (1969), and the sedimentary evolution of the jurassic successions was discussed by sellwood (1972), rolle et al. (1979) and gravesen et al. (1982); the last authors also erected the first formal lithostratigraphic scheme. palynological studies have been undertaken by hoelstad (1985), koppelhus (1991), koppelhus & batten (1992), batten et al. (1994) and koppelhus & nielsen (1994). surlyk & noe-nygaard (1986) interpreted the depositional environments of the pliensbachian hasle formation, and the rich ammonite fauna is described by donovan & surlyk (2003, this volume). surlyk et al. (1995) provided a detailed sequence stratigraphic interpretation of the hettangian– sinemurian sose bugt member (rønne formation). koppelhus & nielsen (1994) and petersen & nielsen (1995) interpreted the depositional environments of the upper pliensbachian – bathonian sorthat and bagå formations. the coal beds, burial depths and amount of uplift based on coal maturation were studied by petersen (1993) and petersen et al. (2003a, b, this volume). jurassic sediments in the subsurface are known from a large number of deep wells drilled for hydrocarbon contents regional setting 149 lithostratigraphy 154 bornholm 155 rønne formation 155 hasle formation 161 sorthat formation 161 bagå formation 166 danish basin 168 gassum formation 168 fjerritslev formation 170 haldager sand formation 173 flyvbjerg formation 174 børglum formation 174 frederikshavn formation 175 danish central graben 176 fjerritslev formation 176 bryne formation 177 lulu formation 180 middle graben formation 183 lola formation 185 heno formation 185 gert member 189 ravn member 190 farsund formation 191 bo member 195 poul formation 197 stratigraphic evolution 197 early jurassic marine deposition 197 middle jurassic uplift 204 middle jurassic fluvial and paralic deposition 204 late jurassic marine deepening 208 late jurassic graben development and sedimentation 209 acknowledgements 211 references 212 and geothermal exploration with supplementary information from seismic surveys. the presence of jurassic rocks from the subsurface of onshore denmark was described for the first time by nørvang (1946) from a borehole situated in northern jylland near the village of vejrum. the vinding-1 well, drilled in 1947, was the first deep onshore well to penetrate the jurassic. gregersen & sorgenfrei (1951) presented a stratigraphic review of two well sections, gassum-1 and vinding-1. a lithological and stratigraphical summary of the jurassic in a large number of wells was given by sorgenfrei & buch (1964). a lithostratigraphic subdivision of the entire jurassic succession was undertaken by larsen (1966), and later refined and partly revised by bertelsen (1978) and michelsen (1978a, 1989a). larsen et al. (1968) described jurassic strata encountered in wells drilled in the strait between denmark and sweden (øresund); petersen (1994) studied the genesis of the coals from these wells. several biostratigraphic studies have been published including those of nørvang (1957) who examined the early jurassic foraminiferal faunas, michelsen (1975) who established a zonation of the lower jurassic based on the ostracod faunas, dybkjær (1988, 1991) and poulsen (1992, 1994, 1996) who established a palynological zonation of the jurassic in the danish basin. the depositional environments and bivalve assemblages of the lower jurassic fjerritslev formation were interpreted by pedersen (1985, 1986). koch (1983) discussed the depositional environments of the middle–upper jurassic sandstone reservoirs, thomsen et al. (1987) described the hydrocarbon plays, and nielsen (1995, 2003, this volume) has described the upper triassic – jurassic depositional sequences and the development of the danish basin and fennoscandian border zone. in the danish north sea sector, the jurassic was first penetrated by the a-2 well, drilled in 1967, and sorgenfrei (1969) indicated the presence of jurassic marine offshore deposits in the north sea. the first lithological and stratigraphic summaries of drilled jurassic successions were presented by rasmussen (1974, 1978). a regional description and a preliminary lithostratigraphic subdivision of the jurassic succession in the central graben were presented by koch et al. (1982). the lithostratigraphic subdivision defined in the danish onshore well sections also applies to the easternmost north sea wells located in the norwegian–danish basin (michelsen 1978a). a lithostratigraphic scheme for the deposits in the central graben was established by jensen et al. (1986), and jurassic basin development was discussed by michelsen et al. (1987). the depositional environments of the middle jurassic bryne and lulu formations and the upper jurassic heno formation were interpreted by johannessen (1995, 1997, 2003, this volume), johannessen et al. (1996), petersen & andsbjerg (1996) andsbjerg (1997, 2003, this volume) and andsbjerg & dybkjær (2003, this volume). biostratigraphic data from the offshore well sections were presented by rasmussen (1974, 1978). the early jurassic ostracod faunas of the o-1 well were described by michelsen (1978b), hoelstad (1986a) studied the palynology of the middle jurassic in the u-1 well, and poulsen (1986, 1991) presented a dinoflagellate cyst biostratigraphy of the middle and upper jurassic in two wells from the central graben. dybkjær (1998) and andsbjerg & dybkjær (2003, this volume) presented palynological data from the middle and upper jurassic. the hydrocarbon plays were described by damtoft et al. (1992); further details on source rocks were provided by petersen et al. (1996, 1998, 2000), petersen & rosenberg (1998) and ineson et al. (2003, this volume). the primary aim of the present paper is to present a complete and updated version of the lithostratigraphic scheme of the entire danish onshore and offshore area (fig. 2). in addition, an account is given of the stratigraphic development during the jurassic period in the danish part of the north sea basin including the danish basin. biostratigraphic and sequence stratigraphic aspects are dealt with in other papers in this volume and will be referred to where relevant. a summary of the distribution and thickness of the lithostratigraphic units penetrated in well sections can be found in nielsen & japsen (1991). regional setting the danish area belongs to the eastern part of the north sea basin, which comprises a number of faultbounded basins separated by structural highs (figs 1, 3). the danish jurassic deposits represent a wide range of environments, from dominantly paralic and coastal in the east to deep marine in the west. differential movements of individual structural features have controlled the distribution of jurassic deposits. the most important structural features are therefore described briefly below. the north sea basin is separated from the fennoscandian shield to the east by the strongly block-faulted fennoscandian border zone (sorgenfrei & buch 1964; baartman & christensen 1975). the zone runs from northernmost jylland, south-eastwards through skåne to the island of bornholm in the baltic sea. it comprises 149 150 egersund subbasin southern vestland arch fiskebank basin north sea baltic sea fennoscandian border zone 10°e8°e 12°e 14°e 10°e8°e 12°e 14°e 16°e 58°n 57°n 56°n 55°n 58°n 57°n 56°n 55°n 6°e4°ea sweden the netherlands germany norway denmark uk skåne bornholm horn graben horsens-1 rønde-1 terne-1 gassum-1 mejrup-1 rødding-1 hyllebjerg-1 vedsted-1 flyvbjerg-1 vinding-1 nøvling-1 oddesund-1 skive-1, -2 kvols-1 børglum-1 haldager-1 frederikshavn-1, -2 hans-1 pernille-1 stina-1 fjerritslev-1,-2 f-1 inez-1 k-1 felicia-1 j-1 well fault high well fault 100 km structural elements of southern scandinavia b national boundaries sorgenfrei–tornquist zone børglum fault haldager fault fjerritslev fault danish basin skurup high ringkøbing– fyn high skagerrak– kattegat platform rø nn e g ra be n bornholm 50 km danish basin central graben mainly marine and paralic jurassic deposits, and was repeatedly overstepped during the jurassic. in the kattegat area and northernmost jylland, the zone includes two different structural elements, the sorgenfrei–tornquist zone and the skagerrak–kattegat platform (figs 1b, 3; eugeno-s working group 1988). the sorgenfrei–tornquist zone was established during late carboniferous – early permian times as a rift zone in conjunction with the oslo graben (liboriussen et al. 1987; eugeno-s working group 1988; ro et al. 1990; michelsen & nielsen 1991, 1993; mogensen & korstgård 1993, 2003, this volume; mogensen 1994, 1996; vejbæk 1997). the zone was periodically active during the mesozoic. block faulting and increased subsidence rates prevailed in the late triassic – early jurassic and during latest jurassic – earliest cretaceous times, while inversion tectonism characterised late cretaceous – early cenozoic times. slow subsidence occurred in 151 ■ ■ ■ ■ gertrud plateau/grabengert ridge late jurassic structural elements normal fault reverse fault salt structures well ▲ ■ 25 km ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ▲ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ▲ ▲ ■ ■ ■ ■ ■ ▲ ▲ ▲ ■ ■ ■ ■ ■ ■ ■ ugle-1 cleo-1 amalie-1 elin-1 west lulu-1 west lulu-3 gert-1 gert-4 gert-2 kim-1 tordenskjold-1 p-1 2/8-3 jeppe-1 gwen-2 t-1 ravn-1 edna-1 ravn-2 w-1 elly-2 elly-3 ringkøbing–fyn high east north sea high deep gorm-1 g-1 m-8 a-2 o-1 e-1 l-1 u-1 lulu-1 nora-1 poul plateau v-1 bo-1 mid north sea high feda graben inge high søgne basin m ads h igh tail end g raben r osa basin coffee soil fault heno plateau baron-2 4°e 56°n salt dome province mandal high c fig. 1. a: outline structure map showing the three main mesozoic tectonic elements: the central graben, the danish basin and the fennoscandian border zone. the red line parallel to the denmark–norway border indicates the location of the geosection in fig. 3. b: simplified structural map of the danish basin with well locations. modified from nielsen (2003, this volume). c: map of the danish central graben showing the late jurassic structural elements and the location of wells mentioned in the text. modified from japsen et al. (2003, this volume). 152 oddesund fm k åg er öd fo rm at io n system series stage ju ra ss ic t ri as si c åsgard formation leek member bo member heno fm lola formation bryne formation danish basin sw ne vedsted formation børglum formation flyvbjerg formation haldager sand formation fj er ri ts le v fo rm at io n gassum formation vinding formation skagerrak formation jydegård formation robbedale formation rabekke formation bagå formation sorthat formation hasle formation galgeløkke mb sose bugt mb r øn ne f or m at io n f-iv mb f-iii mb f-ii mb f-ib f-i mb f-ia fennoscandian border zone risebæk mb munkerup mb frederikshavn formation danish central graben ryazanian volgian kimmeridgian oxfordian callovian bathonian bajocian aalenian toarcian pliensbachian sinemurian hettangian rhaetian norian winterton formation c re ta ce ou s valanginian fjerritslev formation vyl fm poul fmfarsund formation lo w er u pp er m id dl e lo w er u pp er u l u l l u u m l l m u m l u l u u l l u m l u u l u m l annero fm a nn er o fm m ar ie da l f m vilhelmsfält fm röddinge fm ? ? ? paralic and non-marine sandstones, siltstones, mudstones and coals marine mudstones and siltstones unconformity offshore organic-rich marine shales submarine fan sandstones and siltstones shallow marine sandstones and siltstones hiatus fyledalen clay nytorp sand fortuna marl glass sand mb fuglunda mb vitabäck clay helsingborg mb döshult mb pankarp mb katslösa mb rydebäck mb r ya f m h ög an äs f m bjuv mb vallåkra mb h öö r ss t. kågeröd fm skåne bornholmnw se ?? middle graben formation lulu formation fig. 2. lithostratigraphic scheme of the jurassic (including the uppermost triassic and lowermost cretaceous) of the danish central graben, the danish basin and the fennoscandian border zone (bornholm and skåne, southern sweden). colours (jurassic units primarily) indicate overall depositional environments and facies. 153 0. 0 1. 0 2. 0 3. 0 4. 0 5. 0 6. 0 0. 0 1. 0 2. 0 3. 0 4. 0 5. 0 j1 k -1 f1 p1 t1 c en tr al g ra be n sw sw n e sk ag er ra k– k at te ga t pl at fo rm n e 25 k m to rd en sk jo ld -1n or w eg ia n– d an is h ba si n so rg en fr ei –t or nq ui st z on e n or w eg ia n– d an is h ba si n twt sec twt sec c en oz oi c u pp er c re ta ce ou s lo w er c re ta ce ou s u pp er ju ra ss ic lo w er –m id dl e ju ra ss ic t ri as si c z ec hs te in r ot lie ge nd es fi g. 3 . g eo se ct io n f ro m t h e c en tr al g ra b en i n t h e w es t to t h e sk ag er ra k– k at te ga t p la tf o rm i n t h e ea st ; p o si tio n i n d ic at ed i n f ig . 1a . m o d if ie d f ro m v ej b æ k (1 99 7) , b as ed o n se is m ic l in e r t d -8 122 . middle jurassic time (nielsen 1995, 2003, this volume; andsbjerg et al. 2001). the skagerrak–kattegat platform was a stable area that was transgressed during the jurassic and is characterised by eastwards thinning of the relatively undisturbed mesozoic succession. the danish basin, constituting the eastern part of the norwegian–danish basin, is situated south-west of the sorgenfrei–tornquist zone. the basin is bordered to the south by the ese–wnw-trending ringkøbing–fyn high, which consists of a series of shallow fault blocks of precambrian basement (sorgenfrei & buch 1964). the high is of pre-permian age and is separated from the mid north sea high by the north–south-oriented central graben (fig. 1). the ringkøbing–fyn high acted as a submarine intra-basinal high during the early jurassic. middle jurassic uplift of the central north sea affected the ringkøbing–fyn high, which was emergent throughout middle and late jurassic times (michelsen 1978a; koch 1983; ziegler 1990; nielsen 2003, this volume). the high separates the norwegian–danish basin in the north from the north german basin in the south. the danish basin, formerly described as the danish embayment (sorgenfrei & buch 1964), and its extension to the west, the norwegian–danish basin, began to subside in permian times (vejbæk 1989, 1997). it contains a thick succession of jurassic deposits, generally comprising marine lower jurassic claystones, fluvial and paralic middle jurassic sandstones, and marine upper jurassic claystones and sandy claystones. there seems to have been open marine connections with basinal areas to the south, west and north during the early jurassic (michelsen 1978a). after the mid-jurassic uplift of the ringkøbing–fyn high, the connection to the mid-european basinal areas was mainly restricted to the central graben area in the west. south of the ringkøbing–fyn high, in the northern part of the north german basin, marine deposition probably only occurred prior to the mid-jurassic uplift of the high. subsequent erosion seems to have removed much of the lower jurassic deposits, which in the danish area are only known from wells in the northeastern part of the basin. the central graben is a complex structural feature, including a number of fault-bounded basins, highs and platforms (fig. 1c), which possibly started to subside in the early permian (ziegler 1975; gowers & sæbøe 1985). however, it remains uncertain if there was a connection in the late permian between the northern and southern zechstein basins through the graben (vejbæk 1992). a minor angular unconformity separates the triassic from the lower jurassic, indicating a late triassic tectonic event. the early jurassic was characterised by regional subsidence and relative tectonic quiescence, and the central graben was a part of the larger north sea basin. marine lower jurassic deposits equivalent to those known from the danish basin are present in the southern part of the graben. they were probably widely distributed in the north sea basin prior to middle jurassic uplift and erosion of the central north sea area; the lower jurassic deposits, as preserved today, represent erosional remnants (gowers & sæbøe 1985; michelsen et al. 1992). the main rifting of the central graben took place in middle and late jurassic times. major fault-controlled subsidence occurred in the eastern part of the graben, along the bounding coffee soil fault zone, and fluvial-deltaic middle jurassic sediments were deposited along the fault zone, probably representing the initial syn-rift deposits. subsidence rates increased significantly during the late jurassic, and more than 4000 m of clay-dominated sediments were deposited in the eastern part of the graben. wrenching led to considerable lateral variation in subsidence rates, and major depocentres developed such as the tail end graben and the feda graben (møller 1986; vejbæk 1992; japsen et al. 2003, this volume). the depositional area gradually expanded to the west, and marine upper jurassic deposits cover the entire graben area (michelsen et al. 1987; damtoft et al. 1992; johannessen et al. 1996). lithostratigraphy within the last forty years, a large number of lithostratigraphic units have been defined within the jurassic successions of the danish onshore and offshore areas (larsen 1966; bertelsen 1978; michelsen 1978a, 1989a; gravesen et al. 1982; jensen et al. 1986). an overview of the lithostratigraphic units encountered in wells in the danish onshore and offshore areas was presented by nielsen & japsen (1991). all jurassic formations are reviewed here, some formations are revised, and two new formations and two new members are defined (fig. 2). the overall lithology, depositional environments, distribution, thickness and age are described. the formations of bornholm are primarily known from outcrops, and the definitions thus include sedimentary facies; information on gamma-log characteristics from shallow cored wells are also included to provide a basis for comparison with subsurface and offshore sections (fig. 4). the formations in the danish basin and the central graben are only known from wells, mainly drilled for hydrocarbon exploration, and 154 from reflection seismic data. the successions penetrated by wells are mainly represented by cuttings samples, and petrophysical logs are therefore used in the identification, delimitation and description of the subsurface formations. the descriptions including the lithology and log features (predominantly gamma-ray and sonic logs) are shown on figures with subsurface type sections or reference sections. formal and informal members of certain formations are discussed briefly under the ‘remarks’ given to the respective formations. the formations from bornholm are described first, followed by formations from the danish basin, and the lithostratigraphic section is concluded with the formations from the danish central graben. within each area, the formations are described in ascending stratigraphic order; with the exception of the uppermost formation in each area, only the lower boundary is described for each formation, the upper boundary being described under the succeeding formation. in borehole/well sections, the depths are given in relation to the land surface (bornholm), the well reference level (typically kelly bushing (kb)) and mean sea level (msl). metric (si) units are preferred, but wells drilled in feet are listed as such, relative to the well reference level, to ensure the accuracy of primary borehole data. bornholm the jurassic succession of bornholm is referred to four formations, the rønne, hasle, sorthat and bagå formations (fig. 2), of which the sorthat formation is new and the bagå formation is revised. the four formations are included in the bornholm group of gravesen et al. (1982). the lower cretaceous rabekke formation is not dealt with here, although it may reach down into the uppermost jurassic. rønne formation history. the rønne formation was defined by gravesen et al. (1982). type area. the cliffs on the south-west coast of bornholm between sønderborg and næbbe odde where the upper parts of the formation are exposed (fig. 4; gravesen et al. 1982). type sections are defined for the two upper members only, the sose bugt member and the galgeløkke member (gravesen et al. 1982). reference section. the combined cored section of the galgeløkke-1 and -2 wells is designated as a reference section to complement the discontinuous exposures and to facilitate correlation to other subsurface and off155 rønne– hasle fault block nyker block rønne graben arnager–sose fault block bornholm high 3 7 4 5 2 1 5 km cretaceous lower–middle jurassic sorthat formation and middle jurassic bagå fm lower jurassic hasle formation fault lower jurassic rønne formation upper triassic kågeröd formation lower palaeozoic precambrian crystalline basement n 9 6sorthat næbbe odde rønne sønderborg 8 vellensby fig. 4. geological map of south-west bornholm (see inset) showing the position of type localities, and type and reference sections. localities marked with the red lines indicate extensive coastal exposures. 1, munkerup coastal section, type locality of the munkerup member, rønne formation; 2, sose bugt coastal section, type section of the sose bugt member, rønne formation; 3, galgeløkke coastal section, type section of the galgeløkke member, rønne formation; 4, hasle coastal section, type locality of the hasle formation; 5, korsodde coastal section, reference section of the sorthat and bagå formations; 6, hasle klinkerfabrik clay pit, type section of the bagå formation; 7, position of the galgeløkke-1 and -2 core wells, reference section of the rønne formation; 8, position of the hasle-1 core well, reference section of the hasle formation; 9, position of the levka-1 well, reference section of the hasle formation and type section of the sorthat formation. modified from gravesen et al. (1982). 156 390 400 350 360 370 380 330 340 230 240 250 260 270 280 290 300 310 320 depth in m depth in m ✩ ✩ ✩ ✩ ✩ ✩ ✩ ✩ ✩ ✩ ✩ ✩ ✩ ✩ ✩ ✩ ✩ ✩ ✩ ✩ ✩ ✩ ✩ ✩ ✩ ✩ ✩ ✩ ✩ ✩ ✩ 2clay si sand 4 8 mm 2clay si sand 4 8 mm galgeløkke-1, -2 composite core log reference section: rønne formation 0 225 gr api 0 225 gr api 157 120 130 140 160 170 180 190 200 210 220 ✲✲✲ ✩ ✲✲✲✩ 10 20 30 40 50 60 70 80 90 100 110 ✩ ✩ ✩ ✩ ✩ ✩ ✩ ✩ ? depth in m depth in m 150 2clay si sand 4 8 mm 2clay si sand 4 8 mm 0 225 gr api 0 225 gr api shore well sections. the wells were drilled at the same locality, south of the town of rønne and c. 300 m ene of the coastal cliff at galgeløkke where the type section of the galgeløkke member is exposed (gravesen et al. 1982; nielsen 1995). the cored wells encountered almost 400 m of the rønne formation (3.2–400.7 m b. surface) and provided continuous cores and a gammaray log (fig. 5). the type section of the galgeløkke member overlies the cored well sections with a stratigraphic gap of some tens of metres, as interpreted from map evidence and the dip of the exposed strata (gry 1969; jensen & hamann 1989; nielsen 1995). the base of the rønne formation was not reached by the wells; the upper boundary with the overlying hasle formation was cored in the hasle-1 well (fig. 6). lithology and environment. the rønne formation consists of a wide variety of lithologies dominated by laminated or massive mud, heteroliths of mud and finegrained sand, fineto medium-grained sand, and thin coal beds with rootlets. deposition took place in paralic and marine environments including coastal plains, 158 lithology mudstone siltstone heterolith, mud/sand (50%) sandstone conglomerate pebble/granule lags coal calcareous cement pyritic concretions carbonaceous detritus mudstone chips sedimentary structures erosional surface parallel bedding/lamination planar cross-bedding cross-bedding with mudstone drapes cross-bedding with double mudstone drapes trough cross-bedding low-angle cross-bedding hummocky cross-stratification gutter casts cross-lamination and climbing ripples bimodal current-ripple lamination wave ripples flaser bedding wavy bedding lenticular and silt-streaked bedding disturbed bedding load structures water escape structures synaeresis cracks clast clast clast basement clast clast fragments biogenic structures bioturbation degree of bioturbation chondrites isp. diplocraterion isp. equilibrichnion isp. rhizocorallium isp. skolithos isp. teichichnus isp. thalassinoides isp. horizontal burrows large burrows rootlets fossils bivalve plant wood stem claystone siltstone clayey sandstone silty sandstone sandstone conglomerate scattered gravel clasts coal bed and clasts limestone bands limestone marlstone dolomite bands volcanic rocks cored section legend: sedimentological logs stratigraphic logs preceding pages and above: fig. 5. composite core log of the rønne formation in the galgeløkke-1 and -2 wells, located in the rønne–hasle fault block just south of the town of rønne on the west coast of bornholm, c. 300 m ene of the coastal cliff at galgeløkke and just west of the eastern bounding fault of the rønne graben; reference section for the rønne formation. no formal subdivision into members is attempted for this section although it is probably largely within the sose bugt member. the formation was cored from 400.7–3.2 m b. surface; the base of the formation was not penetrated. the galgeløkke-1 well was terminated at 123.75 m b. surface due to technical problems; coring was resumed from 120 m to 400.7 m in the galgeløkke-2 well drilled 3 m from galgeløkke-1 (nielsen 1987). note that to facilitate comparison of the shallow cored well sections with outcrop sections on bornholm, composite logs are organised with base to the lower left and top to the upper right. this is the reverse of that typically adopted for deep wells (e.g. figs 11, 16, 26). modified from nielsen (1995). the accompanying legend is applicable to all subsequent sedimentological and stratigraphic logs. lakes, fluvial channels, tidal channels, tidal flats, marine shoreface and restricted offshore. log characteristics. the gamma-ray logs from the galgeløkke-1, -2 and hasle-1 wells show a highly variable pattern reflecting the wide variety of lithologies, with sand displaying blocky, coarsening-upwards (i.e. values decrease upwards) and fining-upwards (values increase upwards) trends. mud is indicated by relatively uniform high values, while the common heteroliths are shown by intermediate, variable values. lower boundary. onshore, the formation rests unconformably on triassic or cambrian strata. the lower boundary is defined by a change from red, green and variegated clay and greenish sandstones lacking carbonaceous detritus of the triassic risebæk member (kågeröd formation) or from cambrian quartzitic sandstones to light and dark grey to black clay with abundant carbonaceous detritus and subordinate sand of the rønne formation. offshore, in the pernille-1 and stina-1 wells, the rønne formation (munkerup member) overlies upper triassic strata, probably of rhaetian age, mainly consisting of slightly calcareous to marly brownish claystones; the boundary is tentatively placed at the change to non-calcareous claystones with organic detritus. distribution and thickness. the rønne formation may be up to 500 m thick in the eastern part of the rønne graben near the town of rønne (nielsen 1995). the formation extends offshore in the rønne graben as shown by the pernille-1 and stine-1 wells (fig. 1c, see also fig. 31). age. the rønne formation is essentially of hettangian– sinemurian age but may extend into the earliest pliensbachian (gry 1969; koppelhus 1991; surlyk et al. 1995; donovan & surlyk 2003, this volume). the munkerup member is of early hettangian age. the sose bugt member is of late hettangian – late sinemurian or earliest pliensbachian age in the arnager–sose block, while it may be limited to the late hettangian – early sinemurian in the rønne–hasle fault block, where the galgeløkke member probably spans the late sinemurian. subdivisions. gravesen et al. (1982) subdivided the formation, in ascending stratigraphic order, into the munkerup member, the sose bugt member and the galgeløkke member (fig. 2). the munkerup member is known from numerous small clay pits, long since infilled. the description of the member is therefore based on old literature, and a type section has not been measured (gravesen et al. 1982). the member is dominated by grey to black clay with subordinate thin beds of red, brown and white clay, and sandstone. carbonaceous detritus and plant fossils are abundant in the clays, locally forming coal laminae or thin coal beds. the sandstones are very fine-grained, white to yellowish. in addition to the occurrences at the type locality on the coast at munkerup and at vellensby, where a reference section has been defined in a shallow borehole, the member seems to be present in the pernille-1 and stina-1 wells in the rønne graben (fig. 1c, see also fig. 31; nielsen 1995). in these wells, c. 50 m of mainly grey to olive grey and light grey clay is interbedded with fineto medium-grained sand containing a mixed assemblage of early jurassic and reworked triassic spores and pollen comparable to that of the onshore munkerup member (koppelhus 1991; nielsen 1995). according to gry (1969) the thickness of the munkerup clay onshore is 20 m although the upper boundary is not seen. the sose bugt member is recognised onshore bornholm in both the arnager–sose and rønne–hasle fault blocks, and a type section has been defined from the outcrop at sose bugt (fig. 4; gravesen et al. 1982; surlyk et al. 1995). the member appears to be thickly developed offshore in the rønne graben (nielsen 1995). it is dominated by alternating thin beds of fine-grained, crosslaminated sand, commonly with abundant organic detritus, and thin beds of grey laminated clay or heteroliths of sand and clay. rootlet horizons and thin coal beds are common. fine-grained and mediumto coarse-grained sand units, up to 12 m thick, with cross-lamination and cross-bedding are common in the rønne–hasle fault block as shown by the galgeløkke-2 well (fig. 5). the galgeløkke member is dominated by light grey to white or yellowish, cross-laminated, fine-grained sand or heteroliths showing wavy and flaser-laminated sand and clay, particularly in the lower part. large-scale cross-bedded mediumto coarse-grained sand beds with tidal bundles and mud couplets (sellwood 1972; tyge 1990) dominate the middle part of the member. thin coal beds and carbonaceous clays with rootlets occur, as well as organic detritus. the type section of the member is defined in the rønne–hasle fault block (fig. 4), and the member is also present offshore in the rønne graben (gravesen et al. 1982; nielsen 1995). it has, however, not been possible on the basis of the available data to define the boundary between the sose bugt and galgeløkke members with confidence in the galgeløkke-1, -2, pernille-1 and stina-1 wells. the member is probably not present in the arnager–sose block 159 160 clay si sand 2 4 8 mm depth in m 70 80 90 100 110 120 118.3 130 140 150 r øn ne f or m at io n h as le f or m at io n 0 700 gr api 0 700 gr api depth in m 10 20 30 40 50 60 h as le f or m at io n clay si sand 2 4 8 mm hasle-1, core log reference section: hasle formation gr gr (nielsen 1995; surlyk et al. 1995), as originally supposed (gravesen et al. 1982). hasle formation history. the hasle formation was defined by gravesen et al. (1982). type area. the coastal cliff, south of hasle town, bornholm (fig. 4). a type section has not been measured, but representative sections of the formation may be found in surlyk & noe-nygaard (1986). reference section. the cored hasle-1 well (2.6–155.0 m b. surface), drilled c. 1.1 km south of the hasle harbour (fig. 4), provides continuous cores and a gamma-ray log from the lower–middle part of the hasle formation and illustrates the lower boundary with the rønne formation (fig. 6). the cored levka-1 well drilled between the type area of the hasle formation and the type section of the bagå formation in the hasle klinkerfabrik clay pit, c. 2.5 km south of the hasle harbour and c. 1 km north of the clay pit, provides cores from the middle–upper part of the hasle formation (fig. 7). the two wells are designated as reference sections to facilitate correlation to other subsurface and offshore well-sections; together they illustrate the entire formation with some overlap. lithology and environment. at outcrop, the hasle formation consists of rusty yellow to brownish (dark greenish when unweathered), coarse-grained siltstones and very fine-grained sandstones with swaley to hummocky cross-stratification and marked erosion surfaces draped with small pebble lags. thin poorly sorted gravels, greenish grey and brown clays and clay-ironstones occur locally (gravesen et al. 1982; surlyk & noenygaard 1986). the sandstones are intercalated with fossiliferous clays in the area of the town of rønne and the south coast at stampe å (northern end of the korsodde section, see fig. 4), and a diverse marine fauna including numerous ammonites has been described from the clays (malling & grönwall 1909; malling 1911, 1914, 1920; höhne 1933; donovan & surlyk 2003, this volume). clay beds seem to be absent in the type area to the north as clays were not encountered in the hasle-1 and levka-1 wells which in combination illustrate the entire hasle formation (figs 6, 7). deposition took place in a storm-dominated shoreface to offshore environment (surlyk & noe-nygaard 1986). log characteristics. the gamma-ray log from the hasle-1 well (fig. 6) shows intervals with relatively uniform low readings interrupted by log spikes that locally show extremely high values; such high values in the sandstones are caused by concentrations of uranium and thorium (nielsen 1995). lower boundary. the formation conformably overlies the rønne formation and is marked by a change from largely unconsolidated heterolithic sand and clay with carbonaceous detritus to indurated limonitic brownish (weathered) or dark greenish (unweathered) sandstones or thick clay beds containing marine macrofossils. in the hasle-1 reference section, the base of the hasle formation is defined at a marked pebble-strewn surface interpreted as a ravinement surface capping crossbedded and flaser-laminated sandstones with coal particles of the underlying rønne formation (fig. 6, 118.3 m; nielsen 1995). distribution and thickness. the formation is probably up to 140 m thick in the type area and seems to be present offshore in the stina-1 well in the southern part of the rønne graben in a slightly more fine-grained, clayey facies (see fig. 31; nielsen 1995). age. the formation is of early pliensbachian age (donovan & surlyk 2003, this volume). sorthat formation new formation history. the coal-bearing clays and sands that overlie the hasle formation in the rønne–hasle fault block have 161 facing page: fig. 6. hasle-1 well, located in the rønne–hasle fault block, north of the town of rønne on the west coast of bornholm, just west of the eastern bounding fault of the rønne graben (nielsen 1987); reference section for the hasle formation. the lower c. 37 m of the cored section is referred to the rønne formation; the hasle formation was cored from 118.3–2.6 m b. surface. the upper part of the hasle formation is not represented in hasle-1, but is cored in levka-1 (see fig. 7). note that the scale on the gamma-ray log differs from the galgeløkke-1, -2 and levka-1 wells. modified from nielsen (1995); for legend, see fig. 5. 162 so rt ha t fo rm at io n clay sandsi 2 4 8 mm depth in m levka-1, core log reference section: hasle formation type section: sorthat formation 0 gr 225 gr api depth in m clay sandsi 2 4 8 mm 0 225 gr api gr h as le f or m at io n so rt ha t fo rm at io n 229.20 120 130 140 150 160 170 180 190 200 210 220 225 13.4 20 30 40 50 60 70 80 90 100 110 been referred to as the levka, sorthat and bagå beds in older literature (reviewed by gry 1969). a detailed stratigraphic interpretation of the beds has been difficult to achieve owing to complicated block faulting and folding, scattered outcrops and the absence of marine fossils and distinct marker beds. based on the content of megaspores, gry (1969) referred all three units to the middle jurassic with the levka and sorthat beds being roughly contemporaneous, and the bagå beds being possibly slightly younger. gravesen et al. (1982) defined the bagå formation to include the coal-bearing clays and sands of the levka, sorthat and bagå beds and the coal-bearing strata at korsodde and onsbæk (northern end of the korsodde section as indicated on fig. 4). preliminary palynological investigations of samples from the levka-1 core-well and the korsodde section suggested the presence of upper pliensbachian strata (e.b. koppelhus in: nielsen 1987; koppelhus 1988; nielsen & koppelhus 1989). a reassessment of the megaspore collections of h. gry demonstrated the presence of several megaspore species from the bagå formation (sensu gravesen et al. 1982) that suggest the presence of toarcian–aalenian strata, although their long range precluded a definite dating (koppelhus & batten 1992). a thorough palynological–sedimentological study of all available exposures and cores from the lower–middle jurassic has revealed that the marine lower pliensbachian sandstones of the hasle formation are overlain by a succession referable to the levka and sorthat beds that comprises bioturbated sands, heteroliths and clays with thin coal seams containing relatively diverse brackish-marine dinoflagellate assemblages indicative of the upper pliensbachian, toarcian and possibly lower aalenian (koppelhus & nielsen 1994). these paralic deposits are overlain by a succession equivalent to the bagå beds comprising crevasse and fluvial gravels and sands, lacustrine clays, carbonaceous clays and coals belonging to the upper aalenian? and bajocian–bathonian. these new data confirm that the levka and sorthat beds are lithologically different from the bagå beds. thus the levka and sorthat beds are here included in the new sorthat formation, and the bagå formation is revised to include the bagå beds only. both formations are included in the bornholm group of gravesen et al. (1982). the section measured from the sorthat beds by gravesen et al. (1982, fig. 24) covers part of the new sorthat formation. name. after the locality of sorthat on the west coast of bornholm, between the towns of hasle and rønne, where the formation was previously exposed (fig. 4; gry 1969; gravesen et al. 1982; nielsen 1995). type section. the shallow levka-1 core-well is selected as the type section as no exposures are currently available in the sorthat area (fig. 4); this well was drilled between the type area of the hasle formation and the type section of the bagå formation in the hasle klinkerfabrik clay pit, c. 2.5 km south of the hasle harbour and c. 1 km north of the clay pit. the section from 13.4–164.6 m b. surface is represented by both cores and a gamma-ray log (fig. 7). reference section. an intermittently exposed succession, 130–140 m thick at korsodde on the south-west coast of bornholm is selected as the reference section (fig. 8). lithology and environment. the lithology of the sorthat formation is highly variable. the cores and gamma log from the lower c. 112 m of the formation in the levka-1 well reveal sharp-based, fining-upwards units, 3–14 m thick, consisting of coarse-grained, occasionally pebbly sand, overlain by muddy, carbonaceous, micaceous, fineto medium-grained sand, laminated to homogeneous clay and coal seams with rootlets (fig. 7). most of the cored sand shows parallel lamination with subordinate thin beds with cross-bedding, cross-lamination and flaser lamination. large plant fragments and small quartz pebbles are common. occurring between the fining-upwards units are thinly interbedded sand and clay with rootlets and thin coal seams. marine palynomorphs are not found and the interval is interpreted as having been deposited on a coastal or delta plain with fluvial channels, lakes and swamps (koppelhus & nielsen 1994). the upper c. 40 m of the well-section consist of cross-bedded, cross-laminated, wave-rippled and bioturbated sand and heteroliths with sporadic synaeresis cracks, pyrite nodules, planolites isp. and teichichnus isp. burrows and brackish-marine palynomorphs. fining-upwards successions of sand, clay and coal seams are also represented and are characterised 163 facing page: fig. 7. levka-1 well, located in the rønne–hasle fault block north of the town of rønne on the west coast of bornholm, just west of the eastern bounding fault of the rønne graben (fig. 4; nielsen 1987); type section for the new sorthat formation. the formation was cored from 164.6–13.4 m b. surface. a gammaray log was not recorded from the hasle formation. modified from nielsen (1995); for legend, see fig. 5. 164 clay si sand 2 4 8 mm so rt ha t fo rm at io n so rt ha t fo rm at io n so rt ha t fo rm at io n so rt ha t fo rm at io n h as le f or m at io n 3.2 m not exposed m 92.8 m 32 34 36 37.6 60 72 72.5 76 84.6 86 88 40°s 56° 90 92 8 m not exposed 3 m not exposed 11 m not exposed 21.4 m not exposed 8.4 m not exposed 116 114 112 100 102 104 106 108 110 96 98 95 94.7 m 140 138 136 134 130 132 128 126 122 124 120 118 36°s 48° 22°e 30° 49°s 76° 58°s 68° 57°s 67° 62°s 64° 22 24 26 28 20 16 18 12 14 3.6 0 2 m 28.8 clay si sand 2 4 8 mm clay si sand 2 4 8 mm clay si sand 2 4 8 mm korsodde coastal section reference section: sorthat formation reference section: bagå formation fig. 8. coastal section from south-west bornholm, rønne–hasle fault block, close to the eastern bounding fault of the rønne graben; reference section of the new sorthat formation. the formation overlies the hasle formation, but the formation boundary occurs within an unexposed interval of 8.4 m. modified from koppelhus & nielsen (1994); for legend, see fig. 5. 165 ba gå f or m at io n ba gå f or m at io n ba gå f or m at io n so rt ha t fo rm at io n m 200 198 196 194 192 190 188 186 46°s 66° 40°s 70° 68°e 44° 56°s 74° 60°s 38° 5.2 m not exposed clay si sand 2 4 8 mm clay si sand 2 4 8 mm clay si sand 2 4 8 mm m 186 184 182.4 177.2 176 166 168 170 172 174 162 164 m 160 158 156 154 152 150 148 144 146 140 142 by non-marine palynomorphs. this upper c. 40 m interval is interpreted as having been deposited in lagoons, coastal lakes and fluvial channels with the clean sand at the top probably representing a marine shoreface. the reference section at korsodde is also lithologically highly variable (fig. 8). the lower c. 93 m thick unit consists of mediumto coarse-grained, pebbly, cross-bedded and parallel-laminated sand units, up to 6 m thick. thin, fine-grained and cross-laminated sand beds occur. the sand is locally almost black due to abundant organic debris. interbedded with the sand are laminated to almost homogeneous clay beds and fining-upwards heteroliths overlain by laminated dark clay and coaly beds with rootlets. brackish palynomorphs are present. this part of the section is interpreted to have been deposited in fluvial channels, coastal lakes and lagoons. above a pronounced erosion surface (fig. 8, 105.5 m) is a fining-upwards unit, c. 19 m thick, of cross-bedded, coarseto fine-grained sand overlain by heteroliths and clays capped by a coal seam. some cross-beds show tidal structures such as reactivation surfaces, bundles and mud-drapes that yield brackish-marine dinoflagellates. small burrows are also common. larger burrows, including diplocraterion isp., occur in the wavy to flaser bedded heteroliths. the unit is interpreted as the fill of an estuarine channel. it is overlain by bioturbated heteroliths with dinoflagellates and burrows of diplocraterion isp., teichichnus isp., skolithos isp. and planolites isp. well-sorted, fine-grained sand beds with wave-ripples and swaley cross-stratification also occur. these inferred lagoonal deposits are followed by fluvial sands and lacustrine clays. the uppermost part of the formation in the korsodde section consists of very fineto fine-grained, swaley crossstratified and parallel-laminated yellowish–brown sands and sandstones with thin bioturbated and wave-rippled heterolithic beds deposited in the shoreface to offshore transition zone (koppelhus & nielsen 1994). the coals of the levka-1 and korsodde sections were primarily formed in anoxic, water-saturated inter-channel environments on a coastal plain (petersen et al. 2003, this volume). log characteristics. the gamma-ray log shows a very variable pattern reflecting the interbedded sands, heteroliths, muds and coals (fig. 7). both fining-upwards and coarsening-upwards trends are represented, as testified by gamma log values that increase and decrease upwards, respectively. lower boundary. the lower boundary is defined at the change from dark green (unweathered), yellow to indurated brownish (weathered) limonitic marine sandstones of the hasle formation to unconsolidated light greyish fluvial sands, lacustrine grey to dark grey clays, and coals (figs 7, 8). in the type section, the boundary is marked by a change from laminated and cross-bedded, mediumto coarse-grained beach sand of the uppermost hasle formation to coal-bearing sand and clay with roots. in the korsodde reference section, the boundary with the hasle formation occurs within a non-exposed interval of c. 8.4 m, covered by a concrete construction. distribution and thickness. the formation occurs in the rønne–hasle fault block, where it is up to 200 m thick. the formation is also present offshore in the rønne graben, as indicated by seismic data and the stina-1 well (see fig. 31; hamann 1994; nielsen 1995). age. gry (1969) proposed a middle jurassic age for the sorthat, levka and bagå beds based on the megaspore content. however, a reassessment of the megaspores revealed the possible presence of toarcian–aalenian strata (koppelhus & batten 1992). based on the recognition of several dinoflagellate assemblages supported by spores and pollen, koppelhus & nielsen (1994) proposed that the age of the sorthat formation is late pliensbachian – toarcian, possibly extending up into the early aalenian. bagå formation revised formation history. the bagå formation was defined by gravesen et al. (1982) to include the coal-bearing clays and sands in the rønne–hasle fault block traditionally named the levka, sorthat and bagå beds (reviewed by gry 1969). the coal-bearing strata at korsodde and onsbæk were also included in the bagå formation. as discussed above (sorthat formation), the bagå formation as revised here 166 facing page: fig. 9. hasle klinkerfabrik clay pit, type section of the bagå formation; re-measured by koppelhus & nielsen (1994) to cover the full exposure. modified from koppelhus & nielsen (1994); for legend, see fig. 5. 167 clay sandsi 2 4 8 mm clay sandsi 2 4 8 mm clay sandsi 2 4 8 mm clay sandsi 2 4 8 mm m 24 20 22 16 18 12 10 14 8 6 4 2 0 m 72 70 68 64 66 60 58 62 56 54 52 50 48 m 100 88 90 92 94 96 98 84 82 86 80 78 76 74 72 m 48 46 44 40 42 36 34 38 32 30 28 26 24 ✩ ✩ ✩ ✩ ✩ ✩ hasle klinkerfabrik clay pit type section: bagå formation is adopted in a restricted sense to only include the middle jurassic bagå beds of gry (1969). type section. the succession exposed in the hasle klinkerfabrik clay pit, immediately south of the bagå stream, is selected as the type section in accordance with gravesen et al. (1982), who presented 23 m of section. the exposed section has been re-measured to present the full section of the pit (fig. 9). reference section. the upper c. 65 m of the section exposed at korsodde is proposed as a new reference section (fig. 8), as the former reference section used by gravesen et al. (1982, fig. 24) is now referred to the sorthat formation. lithology and environment. the bagå formation includes thick laminated to homogeneous, grey clay units, up to 10 m thick, dark to black coaly clays with rootlets and coal beds, up to 2.5 m thick, and mediumto finegrained, cross-bedded or poorly laminated sand beds forming units up to 5.5 m thick. poorly sorted, muddy and pebbly sand beds, locally with boulders of weathered granite, occur in the upper part (fig. 9). the lower part of the reference section at korsodde comprises poorly parallel-stratified, coarseto very coarse-grained pebbly sand with boulders of cambrian quartzites and lower palaeozoic kaolinised mudstones (fig. 8). the middle part comprises laminated grey clay and dark highly carbonaceous clay. the upper part consists of poorly sorted fineto medium-grained, poorly laminated sand beds, 1.5–3.0 m thick, interbedded with muddy and carbonaceous sand. deposition took place in lakes and swamps, small crevasse channels and lacustrine deltas, and fluvial channels (gravesen et al. 1982; koppelhus & nielsen 1994; nielsen 1995). the coaly clays and coal beds were primarily deposited in open freshwater mires subjected to some siliciclastic deposition. coal beds in the upper part of the formation were deposited on an alluvial fan with a fluctuating watertable. boundaries. the lower boundary is not exposed in the sorthat area and was not encountered in the cored wells. in the korsodde reference section, the lower boundary is defined at a significant pebble-strewn erosion surface separating swaley to hummocky crossstratified, weakly cemented, yellowish brown, finegrained shoreface sandstones of the upper sorthat formation from overlying coarse-grained, pebbly, light grey to white, largely unconsolidated fluvial sands with pebbles and boulders (fig. 8, 145.2 m). onshore, the upper boundary is an erosion surface defining the present-day land surface or the base of the quaternary. the korsodde section is faulted against the upper cretaceous bavnodde greensand. distribution and thickness. the bagå formation is present in the rønne–hasle fault block, where it is estimated to be more than 190 m thick (gry 1969), and is also present offshore in the rønne graben with a similar thickness as indicated by seismic data (jensen & hamann 1989; hamann 1994). age. the bagå beds were referred to the middle jurassic by gry (1969) based on megaspores. hoelstad (1985) referred the bagå beds as exposed in the hasle klinkerfabrik clay pit to the uppermost toarcian – aalenian and bajocian–bathonian. however, the palynoflora used by hoelstad (1985) for the identification of the uppermost toarcian – aalenian in the bagå beds appear to have a longer range than originally supposed (t. hoelstad, personal communication 1994). based on the content of spores and pollen in the bagå beds in the hasle klinkerfabrik clay pit and the reference section from korsodde, the age of the re-defined bagå formation is middle jurassic, probably including the late aalenian, bajocian and bathonian (koppelhus & nielsen 1994). danish basin the upper triassic – lowermost cretaceous succession of the danish basin is referred to six formations, the gassum, fjerritslev, haldager sand, flyvbjerg, børglum and frederikshavn formations (fig. 2). new reference sections are proposed here for the gassum and fjerritslev formations. gassum formation history. the gassum formation was defined by larsen (1966), and redefined by bertelsen (1978). type section. the gassum-1 well, 5406–4980 ft b. kb, 1590–1460 m b. msl, in the danish basin (larsen 1966; bertelsen 1978). reference section. the hyllebjerg-1 well, 2751–2582 m b. kb, 2723–2554 m b. msl, drilled centrally in the danish basin is here selected as a reference section in 168 order to present a modern log suite as the petrophysical logs available from the gassum-1 well are of poor quality (fig. 10). lithology and environment. the formation consists predominantly of light grey to whitish fineto mediumgrained, and in places coarse-grained sandstones, interbedded with grey and greenish grey heteroliths, dark-coloured claystones and a few thin coal beds. an overall deltaic origin was suggested by larsen (1966) and bertelsen (1978); the sandstones are both of fluvial channel and marine shoreface origin, however, and occur as widespread sheets that represent several progradational events. the intervening claystones are dominantly marine with some of lacustrine and lagoonal origin in the south-eastern and north-eastern parts of the basin (nielsen et al. 1989; hamberg 1994; hamberg & nielsen 2000; nielsen 2003, this volume). log characteristics. the formation is characterised by changes between high and low gamma-ray values and rather constant and high sonic readings, reflecting interbedded sandstones and claystones. thin intervals with low sonic values probably reflect coal beds. the gamma-ray and sp logs show different styles in different locations within the danish basin. in the central area of the basin, the lower part of the formation includes two intervals showing low, decreasing-upwards gamma-ray values (fig. 10), while more blocky motifs occur in the northern part of the basin; in the south, the gamma-ray log shows increasing-upwards trends. above is a thin interval with high gamma-ray values, succeeded by an overall decreasing-upwards trend, but interrupted by gamma spikes attaining high values. the top of the formation is commonly characterised by a blocky gamma-ray low. to the north and north-west, the upper levels of the formation are characterised by rather high gamma-ray values, decreasing slightly upwards. lower boundary. decreasing-upwards gamma-ray values and increasing sonic values reflect the transition from the underlying vinding formation to the gassum formation. the boundary is located at the base of the lowermost significant sandstone bed within this transition. the skagerrak formation – gassum formation boundary, in the northern and eastern parts of the basin, is similarly defined at the base of the first significant sandstone bed overlying claystones of the uppermost skagerrak formation. 169 hyllebjerg-1 reference section: gassum formation reference section: fjerritslev formation c hr on os tr at ig ra ph y fo rm at io n m em be r h.s. m b. msl bø rg l. fl . m.j. u . j ur . fiv fiii fii fi fib fia iia iib iic fj er ri ts le v g as su m tr ia ss ic lo w er ju ra ss ic v i. li th ol og y 0 150 140 40 m b. kb gr api 1894 1913 1922 1977 2100 2000 2400 2600 2700 2205 2293 2502 2751 2582 2554 2723 1866 1894 1949 2177 2265 2474 1885 sonic msec/ft fig. 10. hyllebjerg-1 well, located in the danish basin; new reference section for the gassum formation, 2723–2554 m b. msl, and the fjerritslev formation, 2554–1894 m b. msl. børgl., børglum; fl., flyvbjerg; h.s., haldager sand; vi., vinding; m.j., middle jurassic; u. jur., upper jurassic. modified from michelsen (1989a); for legend, see fig. 5. distribution and thickness. the formation is present in the danish basin and locally in the north german basin. maximum thicknesses of more than 300 m occur in the sorgenfrei–tornquist zone. age. late norian – rhaetian over most of the basin, younging towards the north-eastern margin, where the formation is of hettangian – early sinemurian age (michelsen 1975; bertelsen 1978, 1980; l.h. nielsen, l. hamberg and e.b. koppelhus in: nielsen 1995). remarks. the formation was subdivided informally into three members by bertelsen (1978), but new well sections suggest that they are not consistently identifiable. fjerritslev formation history. the fjerritslev formation was defined by larsen (1966), and amended by michelsen (1978a). type section. the fjerritslev-2 well, 2233–1322 m b. kb, 2225–1314 m b. msl, in the danish basin (fig. 11; larsen 1966; michelsen 1978a). reference section. petrophysical logs are not available from the entire section in the fjerritslev-2 well, and the section in hyllebjerg-1, 2582–1922 m b. kb, 2554–1894 m b. msl, is selected as reference section (fig. 10). this section was described in detail by michelsen (1989a). lithology and environment. the formation consists of a relatively uniform succession of marine, dark grey to black, slightly calcareous claystones, with a varying content of silt and siltstone laminae. clay-ironstone concretions are common. siltstones and fine-grained sandstones form a minor proportion of the succession, being most common on the skagerrak–kattegat platform. deposition took place in a deep offshore to lower shoreface environment (michelsen 1975, 1978a; pedersen 1985; nielsen 2003, this volume). log characteristics. the formation is characterised by rather uniform high gamma-ray values, but subordinate lows are present in three of the five lithostratigraphic units described below (fig. 10; f-ia, f-ib, f-ii, f-iii, f-iv). the log-motifs of the sp curve closely conform to those of the gamma-ray curve (fig. 11). the sonic and resistivity values are higher in the three units characterised by gamma-ray lows than in the two intervening units (figs 10, 11). lower boundary. in the danish basin, the lower boundary is defined at an abrupt change from low to high gamma-ray values, indicating the shift from the sandstones of the gassum formation to the claystones of the fjerritslev formation (fig. 10). in some cases, the boundary is characterised by a more gradual transition from low to high gamma-ray values (see figs 30, 32). distribution and thickness. the formation is present over most of the danish basin, in the southern part of the danish central graben, and locally within the norwegian part of the norwegian–danish basin. the thickness is very variable due to mid-jurassic erosion. the maximum thickness recorded in well sections is 911 m in the type section. seismic data indicate a thickness of more than 1000 m in the fjerritslev trough. age. mainly early jurassic (michelsen 1975, 1989a), although locally extending down into the late rhaetian (dybkjær 1991) and up into the early aalenian (michelsen & nielsen 1991; nielsen 1992, 2003, this volume; poulsen 1996). the lower boundary of the formation is diachronous, younging towards the northeast where it is of early sinemurian age. the upper boundary is erosional except for within the sorgenfrei– tornquist zone. subdivision. michelsen (1978a, 1989a) subdivided the formation into four informal members (f-i – f-iv), the lowermost of which can be subdivided into two units, f-ia and f-ib, over much of the danish basin. the upper rhaetian – lowermost sinemurian f-ia unit consists of uniform claystones and claystones interbedded with siltstones, which are locally carbonatecemented. the unit is characterised by decreasingupwards gamma-ray values and increasing sonic values. the lower boundary coincides with the lower boundary of the formation. the unit is recognised in most wells in the danish basin and in a few wells in the danish central graben. the sinemurian – lower pliensbachian 170 facing page: fig. 11. fjerritslev-2 well, located in the sorgenfrei–tornquist zone; type section for the fjerritslev formation, 2225–1314 m b. msl. m. jur., middle jurassic. modified from michelsen (1978a); for legend, see fig. 5. 171 haldager sand fm g as su m tr ia ss ic m .ju r. li th ol og y 0 150 resistivity ohm m2/m 100 fjerritslev-2 type section: fjerritslev formation fiii fiv fii c hr on os tr at ig ra ph y fo rm at io n m em be r li th ol og y c hr on os tr at ig ra ph y fo rm at io n m em be r 50 100 sp millivolt sp millivolt fj er ri ts le v lo w er ju ra ss ic fia f -i f -i i f -i ib fib f -i ia fiib fiic fj er ri ts le v lo w er ju ra ss ic m b. kb m b. kb m b. mslm b. msl 1399 1322 1678 1400 1500 1600 1700 1314 1391 1670 1800 1857 1900 2000 2100 2200 2233 1849 2225 155 resistivity ohm m2/m f-ib unit is dominated by uniform claystones showing constantly high gamma-ray values. the lower boundary is characterised by a distinct log-break to higher gamma-ray values and lower sonic values. the unit is recognised in most wells in the danish basin and also in the danish central graben. the lower–upper pliensbachian f-ii member differs from the deposits below and above in that it consists of faintly laminated claystones interbedded with siltstones and sandstones. towards the north-east, on the skagerrak–kattegat platform, the member is dominated by fine-grained sandstones. the member was subdivided into three units (f-iia to f-iic) by michelsen (1978a, 1989a) in a few wells located in the central part of the danish basin. the member typically shows decreasing-upwards gamma-ray and increasing sonic values, with a distinct log-break on the sonic curve at the base of the member. the member varies laterally in thickness, but is recognised in most wells in the danish basin. the upper pliensbachian – lower toarcian f-iii member is dominated by uniform claystones, which in places are slightly silty. consistently high gamma-ray values typify the member, and a distinct shift to lower sonic values defines the lower boundary. the member is recognised in most wells in the danish basin, except for the southernmost part. 172 c hr on os tr at ig ra ph y fo rm at io n li th ol og y -65 90 haldager-1 type section: haldager sand formation reference section: flyvbjerg formation m b. msl 1120 1068 1275 h al da ge r sa nd m id dl e ju ra ss ic fl yv bj er g u pp er ju ra ss ic bø rg lu m fj er ri ts le v lo w er ju ra ss ic ft b.kb 4300 4198.5 4100 4000 3900 3800 3693 3600 3520 3500 sp millivolt -10 resistivity ohm m2/m 15 fig. 12. haldager-1 well, located in the sorgenfrei–tornquist zone; type section for the haldager sand formation, 1275–1120 m b. msl; reference section for the flyvbjerg formation, 1120–1068 m b. msl. modified from michelsen (1978a); for legend, see fig. 5. flyvbjerg-1 type section: flyvbjerg formation reference section: haldager sand formation fj er ri ts le v h al da ge r sa nd m id dl e ju ra ss ic l. ju ra ss ic u pp er ju ra ss ic fl yv bj er g bø rg lu m li th ol og y 10 150 -5 10 904 944 998 m b. msl 951 991 1045 1050 1000 m b. kb c hr on os tr at ig ra ph y fo rm at io n resistivity ohm m2/m sp millivolt fig. 13. flyvbjerg-1 well, located in the sorgenfrei–tornquist zone; type section for the flyvbjerg formation, 944–904 m b. msl; reference section for the haldager sand formation, 998–944 m b. msl. modified from michelsen (1978a); for legend, see fig. 5. the upper toarcian – lower aalenian f-iv member consists of uniform claystones with varying silt content, interbedded with sandstones that are locally carbonate-cemented. the proportion of sandstone, locally with organic debris and coal fragments, increases in the fjerritslev trough and on the skagerrak–kattegat platform. the member is characterised by a distinct shift to higher sonic values at the base. the presence of sandstone intercalations is identified by gamma-ray lows, and carbonate-cemented levels are reflected by high sonic peaks. the top of the member coincides with the top of the formation. the member is recognised in the north-eastern part of the danish basin, the sorgenfrei–tornquist zone and the skagerrak–kattegat platform. haldager sand formation history. the haldager formation was defined by larsen (1966), and amended by michelsen (1978a) who subdivided the formation into two members, the lower of which was defined as the haldager sand member. the member was later given the rank of formation, and the haldager formation was abandoned (michelsen 1989a). type section. the haldager-1 well, 4198.5–3693 ft b. kb, 1275–1120 m b. msl in the sorgenfrei–tornquist zone is defined as the type section (fig. 12; larsen 1966; michelsen 1978a). reference section. the flyvbjerg-1 well, 1045–991 m b. kb, 998–944 m b. msl in the sorgenfrei–tornquist zone (fig. 13; michelsen 1978a). lithology and environment. the formation consists of thick beds of light olive-grey, fineto coarse-grained occasionally pebbly sandstones and siltstones interbedded with claystones and thin coal beds. the southwestern part of the basin and the skagerrak–kattegat platform is dominated by sandstones deposited in braided rivers and lacustrine mudstones (koch 1983). shallow marine sandstones, estuarine sandstones and mudstones, and marine mudstones occur in the sorgenfrei–tornquist zone (nielsen 2003, this volume). log characteristics. the dominance of thick sandstones (low gamma-ray values) interbedded with thin claystones (high gamma-ray spikes) yields a characteristic blocky pattern which is replicated on the sp and resistivity curves (figs 12, 13). lower boundary. the lower boundary is defined at an abrupt change from high to low gamma-ray values and the corresponding shift of the sp log, indicating the abrupt change from the claystones of the fjerritslev formation to the sandstones of the haldager sand formation (figs 12, 13). a regional unconformity separates the two formations. distribution and thickness. the formation is distributed in the central and northern part of the danish basin, 173 c hr on os tr at ig ra ph y fo rm at io n li th ol og y 10 12 -5 22 ft b. kb 1665 1800 1900 1700 2000 2104 2130 2200 2302 frederikshavn-1 type section: frederikshavn formation reference section: flyvbjerg formation børglum formation fl yv bj er g ve ds . l. c r. fr ed er ik sh av n u pp er ju ra ss ic – l ow er c re ta ce ou s b. 495 636 689 629 m b. msl m .j. h .s . resistivity ohm m2/m sp millivolt fig. 14. frederikshavn-1 well, located on the skagerrak–kattegat platform; type section for the frederikshavn formation, 629–495 m b. msl; reference section of the flyvbjerg formation, 689–636 m b. msl and børglum formation, 636–629 m b. msl. b., børglum; veds., vedsted; h.s., haldager sand; m.j., middle jurassic; l.cr., lower cretaceous. modified from michelsen (1978a); for legend, see fig. 5. in the sorgenfrei–tornquist zone and on the skagerrak– kattegat platform. maximum thicknesses of more than 150 m are found in the sorgenfrei–tornquist zone (including the type section), and a significant thinning is seen south-west and north-east of this zone. age. the formation is generally referred to the bajocian– bathonian although the age is poorly constrained; in the sorgenfrei–tornquist zone, the formation is probably of aalenian–callovian age although may extend into the early oxfordian (michelsen & nielsen 1991; poulsen 1996; see discussion in nielsen 2003, this volume). flyvbjerg formation history. the haldager formation was defined by larsen (1966). michelsen (1978a) amended and subdivided the formation into two members, the upper of which was defined as the flyvbjerg member. the member was later elevated to formation rank, and the haldager formation was abandoned (michelsen 1989a). type section. the flyvbjerg-1 well, 991–951 m b. kb, 944–904 m b. msl, in the sorgenfrei–tornquist zone (fig. 13; michelsen 1978a). reference sections. the haldager-1 well in the sorgenfrei– tornquist zone, 3693–3520 ft b. kb, 1120–1068 m b. msl, and the frederikshavn-1 well on the skagerrak– kattegat platform, 2302–2130 ft b. kb, 689–636 m b. msl (figs 12, 14; michelsen 1978a). lithology and environment. the basal part of the formation consists of light olive-grey sandstones and siltstones with traces of coals and locally abundant roots. it is overlain by olive-grey and greyish to olive-black claystones, interbedded with fossiliferous, calcareous sandstones. the upper part of the formation consists of fine-grained sandstones, which are highly calcareous and contain glauconite and traces of pyrite. the sediments were mainly deposited in a shallow marine to offshore marine environment. log characteristics. the formation is characterised by relatively high gamma-ray values and corresponding sp values at the base, decreasing to low values at the top, and by variable sonic velocities. the overall decrease upwards in gamma-ray values is interrupted by thin intervals with low values, reflecting interbedded claystones and thin sandstones. in most wells, a thick sandstone bed represented by a blocky gamma-ray low and a corresponding sp value occurs at the top of the formation. lower boundary. the lower boundary is defined at an abrupt change from low to higher gamma-ray values and a corresponding shift of the sp curve, indicating the change from the sandstones of the haldager sand formation to the more fine-grained clay-rich deposits lowermost in the flyvbjerg formation (fig. 13). in some wells, a more gradual increase in gamma-ray values characterises the uppermost haldager sand formation, while the lower part of the flyvbjerg formation shows a few gamma-ray lows representing thin sandstone beds (haldager-1 and børglum-1; fig. 12, see also fig. 30). distribution and thickness. the formation occurs in the central part and along the northern margin of the danish basin, in the sorgenfrei–tornquist zone and on the skagerrak–kattegat platform forming a north-eastwards thickening wedge. maximum thicknesses of over 50 m are found in the sorgenfrei–tornquist zone and on the skagerrak–kattegat platform. age. late early or early middle oxfordian to late kimmeridgian. both the base and top of the flyvbjerg formation are diachronous, younging towards the northeastern margin of the danish basin (poulsen 1996). børglum formation history. the børglum formation was defined by larsen (1966) but was subsequently demoted to the børglum member of the bream formation by michelsen (1978a). the bream formation was abandoned by vollset & doré (1984), and the børglum member was later restored to the rank of formation (michelsen 1989a). type section. the børglum-1 well in the sorgenfrei– tornquist zone, 3323–3213 ft b. kb, 990–957 m b. msl (fig. 15; larsen 1966; michelsen 1978a). reference section. the frederikshavn-1 well on the skagerrak–kattegat platform, 2130–2104 ft b. kb, 636–629 m b. msl (fig. 14; michelsen 1978a). lithology and environment. the formation consists of a relatively uniform succession of olive to blackish grey, slightly calcareous, homogeneous claystones with varying contents of silt, mica and pyrite. the sediments were mainly deposited in an offshore marine environment. 174 log characteristics. the formation is characterised by rather uniform high gamma-ray values and low sonic velocities. the sonic curve is rather constant in the lower part of the formation, but becomes serrated in the upper part. lower boundary. the lower boundary is located at the top of the thick sandstone beds uppermost in the flyvbjerg formation (fig. 14) or slightly above (fig. 15). the boundary is defined by an abrupt shift from high resistivities and high sonic velocities to lower resistivities and sonic velocities above the boundary (see fig. 30). distribution and thickness. the formation is distributed over most of the danish basin. maximum thicknesses, up to 300 m, are found in wells located close to the fjerritslev fault (e.g. j-1, see fig. 32); a pronounced thinning occurs towards the north-east, south and southwest. age. earliest early kimmeridgian – early volgian over most of the basin (poulsen 1996). the base and top of the formation are diachronous; the base youngs towards the north-eastern margin of the basin (poulsen 1996), and towards the west the upper part of the formation is contemporaneous with the frederikshavn formation and extends into the ryazanian (michelsen 1989a). frederikshavn formation history. the frederikshavn and skagen formations were defined by larsen (1966) but were subsequently combined and reduced in rank to the frederikshavn member of the bream formation by michelsen (1978a). the bream formation was abandoned by vollset & doré (1984), and the frederikshavn member was subsequently elevated to the rank of formation (michelsen 1989a). type section. the frederikshavn-1 well, 2104–1665 ft b. kb, 629–495 m b. msl, on the skagerrak–kattegat platform (fig. 14; larsen 1966; michelsen 1978a). reference section. the børglum-1 well, 3213–2478.5 ft b. kb, 957–733 m b. msl, in the sorgenfrei–tornquist zone (fig. 15; michelsen 1978a). lithology and environment. the formation consists of siltstones and fine-grained sandstones, greenish grey, light olive-grey and olive-black, slightly calcareous, in places with micro-lignite. the siltstones and sandstones are interbedded with thin silty, greyish black, and calcareous claystones and form 2–3 generally coarseningupwards units separated by claystone-dominated intervals. glauconite and fragments of bivalves and 175 li th ol og y børglum-1 type section: børglum formation reference section: frederikshavn formation -10 15 fr ed er ik sh av n u pp er ju ra ss ic – l ow er c re ta ce ou s 10 sp millivolt 990 957 733 c hr on os tr at ig ra ph y fo rm at io n m b. mslft b. kb 150 bø rg lu m fl yv bj . ve ds . l. c r. 2600 2478.5 2700 2800 2900 3000 3100 3200 3213 3300 3323 resistivity ohm m2/m fig. 15. børglum-1 well, located in the sorgenfrei–tornquist zone; type section for the børglum formation, 957–990 m b. msl. flybj., flyvbjerg; veds., vedsted; l.cr., lower cretaceous. modified from michelsen (1978a); for legend, see fig. 5. ammonites are common in the lower and upper parts of the formation (sorgenfrei & buch 1964; larsen 1966; birkelund & pedersen 1980). in the sorgenfrei–tornquist zone, sandstones are particularly common in the middle part of the formation. close to the north-eastern basin margin, coal beds a few centimetres thick occur in the upper part of the formation. deposition mainly took place in a shallow marine to offshore environment. non-marine conditions prevailed on the skagerrakkattegat platform and locally in the sorgenfrei–tornquist zone during deposition of parts of the middle coarsening-upwards unit and the lower part of the upper unit (larsen 1966; michelsen 1978a). log characteristics. the 2–3 coarsening-upwards successions of siltstones and sandstones are reflected by repeated decreasing-upwards trends on the gamma-ray curve, and corresponding features on the sp curve (figs 14, 15). the uppermost part of the formation is characterised by a gradual increase in gamma-ray values. the sonic velocities are high and rather constant, except for a slight decrease in values at the two clayey intervals that separate the coarsening-upwards successions (hyllebjerg-1; see fig. 30). boundaries. the lower boundary is defined by a decrease upwards in gamma-ray values and corresponding features of the sp curve, reflecting the transition from the claystones of the underlying børglum formation to the more coarse-grained frederikshavn formation. the boundary may be best defined by a minor shift from low to higher sonic velocities (hyllebjerg-1, see fig. 30). the upper boundary with the claystones of the lower cretaceous vedsted formation is also a gradual transition. an increase upwards in gamma-ray values and a shift to lower sonic velocities characterises the boundary in the central part of the danish basin (e.g. hyllebjerg-1). in the sorgenfrei–tornquist zone, a more abrupt boundary is indicated by the sp and resistivity logs (figs 14, 15). distribution and thickness. the formation occurs in the eastern part of the danish basin. maximum thicknesses of more than 230 m are found in wells from the sorgenfrei–tornquist zone (e.g. haldager-1, terne-1). relatively thick and coarse-grained successions are also found in the børglum-1, frederikshavn-1, rønde-1 and horsens-1 wells. the thinner and more fine-grained successions found in the hyllebjerg-1, skive-1, kvols-1, rødding-1 and oddesund-1 wells, located in the central part of the basin, seem to mark the westernmost extent of the frederikshavn formation. the succession previously assigned to the formation by michelsen (1978a) in the nøvling-1, j-1, f-1 and k-1 wells further to the west is fine-grained, mainly clay-dominated, and is best referred to the børglum formation (michelsen 1989a). age. volgian–ryazanian. the formation is time-equivalent with the upper part of the børglum formation in the central and western parts of the norwegian–danish basin (michelsen 1989a; poulsen 1996). danish central graben the jurassic succession of the danish central graben is referred to seven formations, the fjerritslev, bryne, lulu, middle graben, lola, heno and farsund formations, of which the bryne formation is revised and the lulu formation is new (fig. 2). two new members are defined in the heno formation and the informal ‘hot unit’ in the farsund formation is defined as the bo member. fjerritslev formation history, type and reference sections. the formation was defined from the danish basin (larsen, 1966; michelsen 1978a) and a full description is given above. jensen et al. (1986) defined a reference section (o-1 well) in the danish central graben. lithology and environment. the formation is dominated by dark grey, slightly calcareous claystones, as seen in the danish basin (see above). in contrast to the latter area, however, the lower levels of the formation in the central graben are notably more carbonate-rich, with interbeds of calcareous silty claystone and soft marlstone (jensen et al. 1986). log characteristics. uniform high gamma-ray values and a uniform sonic velocity pattern characterise the formation; a more erratic sonic log in the lower levels reflects the presence of calcite-cemented silty claystones and thin soft marlstone beds. lower boundary. in the danish central graben, the formation overlies the triassic clay-dominated winterton formation, which is characterised by very low sonic 176 velocities and high gamma-ray values. the boundary is defined by a marked change from high to lower gamma-ray values and very low to higher sonic values. distribution and thickness. the formation has only been found in wells in the southern part of the danish central graben, where it forms an erosional remnant due to extensive mid-jurassic erosion. the thickness of the formation is thus very variable; the maximum measured thickness of 257 m occurs in the edna-1 well. age. in the m-8, o-1 and u-1 wells, the formation is dated to the hettangian–sinemurian (michelsen 1978a); in the edna-1 and deep gorm-1 wells, the upper part of the formation may extend up into the lower pliensbachian (andsbjerg & dybkjær 2003, this volume). bryne formation revised formation history. the bryne formation was defined in the norwegian central graben (vollset & doré 1984) and was described from the danish central graben by jensen et al. (1986). the deposits in the southern danish central graben referred to the lower graben sand formation by jensen et al. (1986) are here included in the bryne formation. the upper part of the bryne formation as adopted by jensen et al. (1986) is here included in the new lulu formation. with this revision, usage of the bryne formation is in accordance with the original definition of the bryne formation by vollset & doré (1984). type section. the norwegian 9/4-3 well, 2613–2507.5 m b. kb (vollset & doré 1984, fig. 29). reference section. the danish lulu-1 well, 12134–11712 ft b. kb, 3669–3540 m b. msl, was selected as the danish reference section by jensen et al. (1986). however, new correlations show that much of the lower part of the formation is missing in lulu-1, probably due to the location of the well on a salt structure that was active in jurassic times. the more complete succession in the danish west lulu-1 well, 12602–11795 ft b. kb, 3801– 3555 m b. msl, is selected as the new reference section (fig. 16). lithology and environment. the bryne formation is characterised by laterally extensive sandstone units separated by thick successions of interbedded sandstones and mudstones (andsbjerg 1997). the laterally extensive sandstone bodies vary from fining-upwards or blocky units, 2–10 m thick, to more than 20 m thick amalgamated units (e.g. uppermost 20 m of bryne formation, fig. 16b). the sandstones are light to medium grey and yellowish brown, very fineto medium-grained. they are characterised by trough and planar cross-bedding, and ripple cross-lamination. contorted bedding and other soft sediment deformation structures are also common. associations of sedimentary structures that may indicate tidal influence occur in the uppermost laterally extensive sandstones, suggesting an origin as estuarine channel sands. sandstones in the lower and middle parts of the formation were deposited in fluvial channels (andsbjerg 2003, this volume). the successions that separate major sandstone units are fining-upwards or finingto coarsening-upwards and commonly 20–60 m thick. they consist of grey to brown very fineto medium-grained sandstones interbedded with siltstones, claystones and occasional thin coals. the sandstone beds are generally less than 2 m thick and show crossbedding, current ripple cross-lamination, climbing ripple cross-lamination, parallel lamination and chaotic bedding. dark to light grey mudstone and siltstone beds may be up to 5 m thick, and most commonly are structureless or show parallel lamination. the finegrained deposits represent a variety of overbank environments ranging from crevasse splays and crevasse channels through distal levee to ponds and lakes. abundant root traces, thin coals and strongly mottled siltand mudstones testify to vegetated substrates and soil development. log characteristics. most of the lithologies are easily distinguished on gamma-ray logs. however, coals are best recognised on the sonic log as distinct low-velocity peaks while high-velocity peaks reflect carbonate-cemented 177 following two pages: fig. 16. a: west lulu-1 well, located in the søgne basin, danish central graben; new reference section for the bryne formation, 3801–3555 m b. msl. cored section indicated by solid bar and illustrated in b. note that the boundary between the lulu and lola formations in this well is marked by the change from a coal bed to marine mudstones and is thought to be a faulted contact (fig. 16b; andsbjerg 2003, this volume). for legend, see fig. 5. 178 m id dl e ju ra ss ic u pp er ju ra ss ic br yn e lu lu lo la br yn e m id dl e ju ra ss ic tr ia ss ic west lulu-1 reference section: bryne formation 0 180 0 180 160 0160 0 11600 11687 11700 3523 11795 11800 11900 12000 12100 3555 c hr on os tr at . fo rm at io n 3801 ft b. kb m b. msl ft b. kb m b. msl c hr on os tr at . fo rm at io n 12200 12300 12400 12500 12602 12600 gr api sonic msec/ft gr api sonic msec/ft a 179 br yn e fo rm at io n m id dl e ju ra ss ic br yn e fo rm at io n lo la f m lu lu f or m at io n 0 180 west lulu-1 0 180 normal fault siclay sand gr siclay sand gr gr api gr api m id dl e ju ra ss ic 11800 11900 11795 11700 ft b. kb 12000 12100 ft b. kb b deposits. the gamma-ray logs typically show low values for the sandstones. channel sandstones may show gamma-ray values that increase upwards, indicative of fining-upwards units (e.g. fig. 16b, 12035–12016 ft b. kb), or box-like, uniformly low values (e.g. fig. 16b, 11840–11797 ft b. kb). the intermediate to high gammaray values of floodplain deposits commonly show increasing-upwards (e.g. fig. 16a, 12512–12437 ft b. kb) or increasing-upwards to decreasing-upwards (e.g. fig. 16a, 12200–12165 ft b. kb) gamma-ray patterns. lower boundary. the bryne formation unconformably overlies claystones of the lower jurassic fjerritslev formation in the southern part and triassic and permian rocks in the northern part of the danish central graben. the boundary is marked by a distinct increase in gammaray values where the bryne formation overlies permian salt, whereas the boundary elsewhere may vary from highly distinct to indistinct depending upon the underlying lithology. distribution and thickness. the formation is found in the søgne basin, in the tail end graben, in the salt dome province and along the south-eastern fringe of the heno plateau. the thickness of the bryne formation is very variable, varying from a few tens of metres in marginal parts of the basin to commonly more than 200 m. the thickest section, 289 m, was recorded in the amalie-1 well in the southernmost søgne basin (base not penetrated). age. the relatively sparse palynological data suggest that the lower part of the bryne formation is aalenian or earliest bajocian in age. based on the occurrence of the dinoflagellate cyst impletosphaeridium varispinosum in the uppermost estuarine sandstones of the bryne formation in the west lulu-1 and west lulu-3 wells, the top of the bryne formation is dated to the latest bathonian – earliest callovian (andsbjerg & dybkjær 2003, this volume). remarks. the middle jurassic deposits of the søgne basin in the northern part of the danish central graben were previously referred to the bryne formation (jensen et al. 1986). the uppermost part of the succession consists of interfingering shallow marine and paralic sandstones and mudstones; this uppermost unit differs from the remainder of the middle jurassic succession both in terms of lithology and depositional environment. in accordance with the original definition of the bryne formation (vollset & doré 1984), these deposits are no longer assigned to the bryne formation but are referred to the new lulu formation (defined below), equivalent to the sandnes formation in the norwegian sector. the middle jurassic succession in the southern part of the danish central graben was referred to the central graben group (the central graben subgroup of van adrichem boogaert & kouwe 1994) by jensen et al. (1986) as well control was missing in the central part of the graben. however, with more well data available it has become clear that the lower part of the middle jurassic succession in the southern danish central graben shows greater affinity to the bryne formation than to the lower graben formation as originally defined by nam & rgd (1980) and modified by van adrichem boogaert & kouwe (1994). in core sections, it is evident that with respect to both lithology and depositional environment this part of the succession cannot be distinguished from that of the bryne formation of the northern danish central graben. it is thus proposed to include the middle jurassic deposits from the southern central graben in the bryne formation. lulu formation new formation history. the succession that makes up the new lulu formation was included in the upper part of the bryne formation by jensen et al. (1986). vollset & doré (1984) defined the equivalent succession in the norwegian sector, penetrated in the fiskebank basin and the egersund subbasin, as a separate formation, the sandnes formation. this unit is missing from most of the more than 100 km wide area that separates the southern part of the søgne basin from the central fiskebank basin. it is thus preferred not to adopt the term sandnes formation for the unit in the southern søgne basin and northern tail end graben. name. from the lulu structure in the søgne basin which was the target for the lulu-1 well (fig. 1c). type section. the danish lulu-1 well, 11813–11715 ft b. kb, 3571–3541 m b. msl (fig. 17). reference section. the danish west lulu-3 well, 12035– 11865 ft b. kb, 3633–3581 m b. msl (fig. 18). lithology and environment. the lulu formation is characterised by coarsening-upwards wedges of shallow 180 marine sandstones, which interfinger with paralic sandstone successions towards the western basin margin. laterally extensive coaly claystones with coal seams, up to 5 m thick, are characteristic of the formation, in particular the lower part (e.g. fig. 18, west lulu-3, 12035–12009 ft b. kb). the shallow marine wedges in the deeper parts of the søgne basin and northern tail end graben each show an 8–12 m thick coarseningupwards succession (fig. 17, lulu-1, 11794–11773 and 11757–11729 ft b. kb). these successions are dominated by hummocky cross-stratified sandstones with interbeds of mudstone and siltstone, overlain by gradually thicker and more coarse-grained sandstone beds showing swaley cross-stratification and planar crossbedding. the sandstones are light grey to greyish brown and most commonly well-sorted, very fineto 181 ft b. kb ft b. kb m b. msl c hr on os tr at ig ra ph y fo rm at io n li th ol og y lulu-1 type section: lulu formation clay si sand gr 160 sonic msec/ft 40 u .j. lu lu lo la br yn e m id dl e ju ra ss ic 0 gr api 125 0 gr api 180 11700 11715 11800 11813 11900 11750 11800 11850 3541 3571 fig. 17. lulu-1 well, located in the søgne basin, danish central graben; type section for the new lulu formation, 3571–3541 m b. msl. cored section indicated by solid bar and illustrated in the core log. u.j., upper jurassic. for legend, see fig. 5. 182 ft b. kb m b. msl ft b. kb m id dl e ju ra ss ic br yn e lu lu lo la clay si sand gr c hr on os tr at ig ra ph y fo rm at io n gr api gr api sonic msec/ft li th ol og y west lulu-3 reference section: lulu formation 0 150 160 40 0 75 3633 358111865 11900 11800 12000 12035 12100 11900 12000 fig. 18. west lulu-3, located in the søgne basin, danish central graben; reference section for the new lulu formation, 3633–3581 m b. msl. cored section indicated by solid bar and illustrated in the core log. for legend, see fig. 5. fine-grained. the coarsening-upwards successions are interpreted as prograding shoreface deposits. along the western margin of the søgne basin and tail end graben, the succession is dominated by backbarrier deposits, consisting mainly of sandstones with evidence for strong tidal influence. the sandstones are commonly moderately to well-sorted, dark and brownish grey to light brown, and show cross-bedding, current and climbing ripple-lamination and abundant double mud-drapes and bundled foresets. bioturbated intervals occur commonly (fig. 18, 11980 ft b. kb in west lulu-3). only thin units of shallow marine sandstone and mudstone are present in the paralic wedges (fig. 18, 11904 ft b. kb. in west lulu-3; andsbjerg 2003, this volume). the coal seams, 0.1 – > 1 m thick with a cumulative thickness of up to 5 m, were deposited on a coastal plain during relative sea-level rise (petersen & andsbjerg 1996). log characteristics. the heterogeneous lithologies are reflected by variable log patterns. two 8–12 m thick units showing gamma-ray values that decrease upwards in the deeper parts of the søgne basin and tail end graben represent coarsening-upwards successions of shoreface and beach sandstones. in the western parts of the søgne basin and tail end graben, 0.5–5 m thick units show gamma-ray values that are consistently low or increase slightly upwards; these units represent channel fills. units with high or decreasing-upwards gamma-ray values between the channel-fills represent lagoonal, tidal flat and bay-head delta environments (andsbjerg 2003, this volume). sonic log intervals 0.5–5 m thick that show very low sonic velocities (usually corresponding to low gamma-ray readings) represent coal deposits. lower boundary. the formation conformably overlies the bryne formation over most of the danish central graben. the lower boundary is placed at the base of the lowermost laterally persistent major coal bed, or at the base of a correlative unit of interbedded sandstones and organic-rich mudstones with thin coals, as seen in amalie-1 and lulu-1. distribution and thickness. the lulu formation is widely distributed in the søgne basin and the northern tail end graben. in the southern part of the tail end graben and in the northern part of the salt dome province, the lulu formation is replaced by the partly time-equivalent, mudstone-dominated middle graben formation (nam & rgd 1980; jensen et al. 1986; van adrichem boogaert & kouwe 1994). the thickness of the lulu formation varies from 30–60 m. age. the formation is of early–late callovian age. this is indicated by the occurrence of the latest bathonian – earliest callovian dinoflagellate cyst impletosphaeridium varispinosum in the uppermost estuarine sandstones of the bryne formation in the west lulu-3 well and the appearance of the late callovian – middle oxfordian dinoflagellate cyst liesbergia scarburghensis in the basal part of the lola formation in several wells (andsbjerg & dybkjær 2003, this volume). middle graben formation history. the middle graben shale formation was defined in the dutch north sea sector by nam & rgd (1980), and extended to the danish central graben by jensen et al. (1986). the name was modified to the middle graben formation by van adrichem boogaert & kouwe (1994). 183 c hr on os tr at ig ra ph y fo rm at io n gr api sonic msec/ft li th ol og y m b. mslft b. kb u-1 reference section: middle graben formation 0 180 180 40 10853 10668 10700 10800c al lo vi an m id dl e g ra be n br yn e lo la c al lo v. (? ) 3223 3279 fig. 19. u-1 well, located in the southern part of the danish central graben; reference section for the middle graben formation, 3279–3223 m b. msl. callov., callovian. modified from jensen et al. (1986); for legend, see fig. 5. type section. the dutch f03-03 well, 3090–2670 m b. kb (nam & rgd 1980). reference section. the u-1 well, 10853–10668 ft b. kb, 3279–3223 m b. msl, is used as a reference section in the danish sector (fig. 19; jensen et al. 1986). lithology and environment. the formation consists of dark brownish grey claystones interbedded with siltstones and rare sandstone beds. coal beds, locally several metres thick, may occur in the lower part of the formation. deposition took place in swampy environments, probably interdistributary bays or coastal lagoons. log characteristics. the gamma-ray log commonly shows a response slightly lower than the overlying lola formation but considerably higher than that of the underlying bryne formation. however, some coal beds may be characterised by a very low gamma-ray response and a low sonic velocity. lower boundary. the middle graben formation in the danish central graben overlies the bryne formation. the boundary is placed at the base of a prominent coal bed, usually the lowermost of a series of coal beds and commonly characterised by a low gamma-ray and a low sonic velocity log response. distribution and thickness. the formation extends from the southern north sea into the salt dome province of the danish central graben. it is not recognised north of the salt dome province, where the bryne formation is overlain by the lulu formation in the danish sector. the thickness of the middle graben formation ranges from 15–56 m in the danish sector. age. based on datings of the underlying deposits belonging to the bryne formation by hoelstad (1986a), the formation is not older than late callovian. there is no good data to constrain the youngest age of the formation, although it probably ranges into the early or middle oxfordian, based on the occurrence of the dinoflagellate cyst liesbergia scarburghensis. the formation may be slightly diachronous, younging southwards to the 184 l ol a c hr on os tr at . fo rm at io n li th ol og y 0 180 180 40 u-1 type section: lola formation ft b. kb 2884 o xf or di an c al lo vi an m . g ra be n 3223 m b. msl 10600 10500 10400 10300 10200 10100 10000 9900 9800 9700 9600 9500 9555 9400 10668 10700 k im m er id gi an fa rs un d gr api sonic msec/ft fig. 20. u-1 well, located in the southern part of the danish central graben; type section for the lola formation, 3223–2884 m b. msl. m. graben, middle graben. modified from jensen et al. (1986); for legend, see fig. 5. dutch north sea area, where it has an early–middle oxfordian age. lola formation history. the lola formation was defined by jensen et al. (1986). type section. the u-1 well, 10668–9555 ft b. kb, 3223–2884 m b. msl, in the danish central graben (fig. 20; jensen et al. 1986). lithology and environment. the formation consists of dark olive-grey to grey claystones, with organic material, mainly of terrestrial origin. deposition took place in a low energy, offshore open marine environment. log characteristics. the formation is characterised by fairly constant, high gamma-ray readings and relatively low sonic velocities. a few high velocity peaks corresponding to dolomite or limestone beds may occur (fig. 20). lower boundary. the formation overlies the lulu formation in the northern part of the danish central graben and the middle graben formation in the southern part of the danish central graben. on structurally high areas along the western margin of the danish central graben and on local highs, the formation rests unconformably on triassic or older strata. the lower boundary with the lulu formation is placed at an abrupt, fairly large increase in gamma-ray values and a minor decrease in sonic velocity (figs 17, 18). the boundary with the middle graben formation shows a sudden and fairly large drop in sonic velocity (figs 19, 20). distribution and thickness. the formation is present in the eastern and southern parts of the danish central graben. it thins west of the tail end graben and probably continues southwards and grades into the lower part of the kimmeridge clay formation in the dutch north sea area (michelsen & wong 1991). maximum thicknesses of c. 1000 m occur in the tail end graben. age. late callovian/early oxfordian – late kimmeridgian. the base of the formation is considered diachronous, younging southwards. the uppermost part is not younger than early kimmeridgian in the elly-2 well in the tail end graben (andsbjerg & dybkjær 2003, this volume). heno formation revised formation history. the heno formation was defined by jensen et al. (1986) to include upper jurassic shallow marine sandstones in the danish central graben. wells drilled in recent years have demonstrated that sandstones of the heno formation interfinger with claystones of the lola formation. in the southern part of the feda graben, two sandstone units are separated by claystones and south and east of the feda graben, on the heno and gertrud plateaus, clayey sandstones separate two discrete sandstone units. the type section of the heno formation, as defined by jensen et al. (1986), comprises only the upper sandstone unit (the new ravn member). the informal terms ‘basal sandstone unit’ and ‘heno equivalent’ have been used for the lower sandstone unit by bergan et al. (1989), michelsen & wong (1991), nielsen & japsen (1991), söderström et al. (1991), 185 c hr on os tr at ig ra ph y fo rm at io n m em be r gr api sonic msec/ftli th ol og y fa rs un d m b. mslft b. kb 13520 13500 13400 13820 13800 13700 13600 13839 4184 4178 4087 w-1 type section: heno formation k im m er id gi an permian h en o r av n lola 0 150 140 40 fig. 21. w-1 well, located on the heno plateau, danish central graben; type section for the heno formation, 4178–4087 m b. msl. modified from jensen et al. (1986); for legend, see fig. 5. 186 c hr on os tr at ig ra ph y fo rm at io n m em be r gr api gr api sonic msec/ft li th ol og y fa rs un d 15500 15400 m b. mslft b. kb ft b. kb 15526 4694 15600 15756 16038 15700 15800 15900 16000 16100 16200 16300 16200 16300 16328 4764 4850 4938 gert-1 reference section: heno formation type section: gert member k im m er id gi an permian lo la h en o g er t h en o r av n 0 150 140 40 clay si sand gr 0 75 fig. 22. gert-1 well, located in the northern part of the danish central graben; new reference section for the heno formation and type section for the new gert member (4938–4850 m b. msl). cored section indicated by solid bar and illustrated by core log. note that this well also encountered the new ravn member (4764–4694 m b. msl) separated from the gert member by 86 m of lola formation mudstones. for legend, see fig. 5. damtoft et al. (1992), michelsen et al. (1992), johannessen & andsbjerg (1993) and johannessen et al. (1996). the revised heno formation, as proposed here, includes both sandstone units that are given the rank of members, the gert and ravn members. type section. the w-1 well, 13820–13520 ft b. kb, 4178–4087 m b. msl, in the central graben, heno plateau, was selected as type section by jensen et al. (1986; fig. 21). reference section. the gert-1 well in the easternmost part of the feda graben bordering the gertrud plateau, and the ravn-1 well on the heno plateau are here selected as reference sections for the heno formation (figs 22, 23). in the gert-1 well, the lower sandstones, 16328–16038 ft b. kb, 4938–4850 m b. msl, are referred to the gert member and the upper sandstones, 15756–15526 ft b. kb, 4764–4694 m b. msl, are referred to the ravn member. claystones of the lola formation lie between the two members. in the ravn-1 well, the sandstone interval 4173–4067 m b. kb, 4132–4026 m b. msl, is referred to the ravn member. lithology and environment. the formation consists of predominantly grey to white very fineto fine-grained sandstones and minor siltstones interbedded with claystones and coal beds. conglomerate beds, 0.08–2 m thick, are present on plateau areas. the gert member was deposited in a back-barrier and marine shoreface environment while 187 c hr on os tr at ig ra ph y fo rm at io n m em be r li th ol og y m b. mslm b. kb m b. kb 4067 4026 4173 4200 4100 4100 4150 4132 gr api gr api sonic msec/ft ravn-1 reference section: heno formation type section: ravn member k im m er id gi an o xf ? r av n fa rs un d lo la h en o 0 150 140 40 500 clay si sand gr fig. 23. ravn-1 well, located on the heno plateau, danish central graben; type section for the new ravn member (heno formation), 4132–4026 m b. msl. cored section indicated by solid bar and illustrated by core log. oxf., oxfordian. for legend, see fig. 5. the ravn member was deposited in a marine shoreface environment (johannessen & andsbjerg 1993; johannessen et al. 1996; johannessen 1997, 2003, this volume). log characteristics. the formation is characterised by low gamma-ray and high sonic readings representing sandstones. some higher gamma-ray and lower sonic readings represent siltstone and claystone beds. lower boundary. in the type section, the formation rests on upper oxfordian – lower kimmeridgian claystones of the lola formation (andsbjerg & dybkjær 2003, this volume). in the gert-1 well, the base of the formation is defined by the base of the lowermost sandstones and claystones of the gert member unconformably overlying lower permian volcanic rocks. 188 c hr on os tr at ig ra ph y fo rm at io n m em be r gr api sonic msec/ftli th ol og y m b. mslm b. kb m b. kb 4940 5000 5014 4902 5041 5003 4976 4970 4980 4950 4960 0 60 gr apijeppe-1 reference section: gert member k im m er id gi an permian r av n g er t fa rs un d h en o 0 150 140 40 clay si sand gr fig. 24. jeppe-1 well, located in the northern part of the danish central graben; reference section for the new gert member (heno formation), 5003–4976 m b. msl. for legend, see fig. 5. distribution and thickness. the formation is present in the feda graben and on the gertrud and heno plateaus in the northern part of the danish central graben. the maximum thickness of the formation on the heno plateau is 106 m in the ravn-1 well. in the feda graben, the total thickness of the heno formation is 158 m in the gert-1 well (gert member 88 m, and ravn member 70 m), where the formation interfingers with 86 m of claystones of the lola formation. age. the formation was assigned a middle–late oxfordian to kimmeridgian age (jensen et al. 1986), but new biostratigraphic data indicate a latest early – late kimmeridgian age (andsbjerg & dybkjær 2003, this volume). subdivision. in the eastern part of the feda graben, the gertrud plateau and the western part of the heno plateau, the lower part of the formation is characterised by a serrated gamma-ray and sonic log pattern including very high gamma-ray peaks corresponding to very low sonic values, recording the presence of thin coal and clay beds within the sand-dominated section. this part of the formation is defined as the new gert member, described below. in the same area, the upper part of the formation is characterised by moderate to low gamma-ray and moderate to high sonic readings representing sandstone facies seen in the type section of the w-1 well. this part of the formation is defined as the new ravn member, described below. the type section of the heno formation in the w-1 well comprises only the ravn member (fig. 21). remarks. other shallow marine sandstones somewhat younger than the heno formation may be present in the western part of the danish area (damtoft et al. 1992). a sandy siltstone succession in the kim-1 well at the eastern margin of the grensen nose basin is thought to be a distal equivalent of the early–middle volgian shallow marine ‘fife sandstones’ in the uk sector (mackertich 1996; andsbjerg & dybkjær 2003, this volume). similar sandstones are expected to be present in the intervening area of the ål and grensen nose basins. gert member new member name. the gert member is named after the gert ridge situated at the eastern margin of the feda graben. type section. the gert-1 well, 16328–16038 ft b. kb, 4938–4850 m b. msl, is proposed as the type section (fig. 22). reference section. the jeppe-1 well, 5041–5014 m b. kb, 5003–4976 m b. msl, on the gertrud plateau, is selected as the reference section (fig. 24). lithology and environment. the member consists predominantly of grey to white very fineto fine-grained sandstones and minor siltstones interbedded with claystones and coal beds (fig. 22). the sediments were deposited in back-barrier and shoreface environments. log characteristics. the lower part of the member in the type section is characterised by serrated gamma-ray and sonic log motifs. log readings from the upper part are more uniform and smooth. the very high gammaray values and very low sonic values in the serrated section represent thin claystone and coal beds interbedded with sandstones. the upper part with low gamma-ray and high sonic values consists of clean sandstones without coal and clay beds. the uppermost part of the member is characterised by a step-wise increase in gamma-ray values and decrease in sonic values, representing a fining-upwards succession from sandstones to the claystones of the overlying lola formation (fig. 22). boundaries. in the type and reference sections, the member rests unconformably on lower permian volcanic rocks. in the gert-1 well, the upper boundary is defined at the change to the clay-dominated lola formation marked by a significant log shift to higher gamma-ray values and lower sonic velocities. in the jeppe-1 well reference section (fig. 24) and the gert-2 well, the gert member is directly overlain by the ravn member, as offshore claystones of the lola formation are absent. in these wells, the top of the gert member is defined at the highest gamma-ray and lowest sonic readings, defining the base of the coarsening-upwards ravn member. distribution and thickness. the member is present in the northern part of the danish central graben with a maximum thickness of 88 m in the gert-1 well at the eastern margin of the feda graben. on the adjacent gertrud plateau, the thickness of the member is reduced to 27 m in the jeppe-1 well and to 23 m in the gert-2 well. age. the member is of latest early – late kimmeridgian age (johannessen et al. 1996; andsbjerg & dybkjær 2003, this volume). 189 ravn member new member name. the member is named after the ravn-1 and ravn-2 wells situated on the south-eastern part of the heno plateau (fig. 1c). type section. the ravn-1 well, 4173–4067 m b. kb, 4132– 4026 m b. msl, is selected as the type section (fig. 23). reference section. the gwen-2 well, 14084–13857 ft b. kb, 4256–4187 m b. msl, on the gertrud plateau, is selected as the reference section (fig. 25). 190 gwen-2 reference section: ravn member k im m er id gi an triassic lo la h en o r av n 0 140 c hr on os tr at ig ra ph y fo rm at io n m em be r li th ol og y ft b. kb ft b. kb 13857 14290 14300 14200 14084 14100 14000 13900 13800 13700 13600 14400 m b. msl 4187 4256 4319 fa rs un d vo lg ia n 14100 14000 clay si sand gr 0 160 140 40 gr api gr api sonic msec/ft fig. 25. gwen-2 well, located on the gertrud plateau, danish central graben; reference section for the new ravn member (heno formation), 4256–4187 m b. msl. modified from jensen et al. (1986); for legend, see fig. 5. lithology and environment. the ravn member consists predominantly of grey to white, very fineto fine-grained sandstones, muddy sandstones and minor siltstones that are bioturbated to such a degree that almost no primary sedimentary structures are preserved (fig. 23). conglomerate beds, 0.08–2 m thick, locally associated with mediumgrained sandstones, are present on plateau areas. in the graben areas, the member is dominated by clayey, very fine-grained sandstones and siltstones. calcareous beds and/or concretions occur. deposition took place in a marine shoreface environment. the conglomerates represent river or beach deposits that were later winnowed during transgressive erosion (johannessen et al. 1996; johannessen 2003, this volume). log characteristics. in the ravn-1 and gwen-2 wells (figs 23, 25), the lower and upper parts of the member include clayey sandstones characterised by relatively high gamma-ray and low sonic readings. the middle part of the member consists of sandstones and conglomerates characterised by low gamma-ray and high sonic readings. the overall pattern is thus characterised by decreasing-upwards gamma-ray values and increasing-upwards sonic values followed by increasing gamma-ray values and decreasing sonic values. in the gert-1 and jeppe-1 wells, the member does not include clayey sandstones in the upper part (figs 22, 24). the gamma-ray and sonic logs are strongly serrated in the ravn-1 well, less serrated in gwen-2, and relatively smooth in gert-1 and jeppe-1 (figs 22–25). this may be due to variation in the development of calcareous beds and/or concretions. lower boundary. the member overlies offshore claystones of the lola formation in the ravn-1 well type section. the base of the member is defined by a shift to lower gamma-ray and higher sonic readings marking the base of an overall coarsening-upwards succession beginning with clayey sandstones (figs 23, 25). in the gwen-2 reference section and the gert-1 well, the member overlies sandy offshore claystones of the lola formation (figs 22, 25), whereas in the jeppe-1 well, the ravn member directly overlies the gert member (fig. 24). distribution and thickness. the member is present in the feda graben and on the gertrud and heno plateaus in the northern part of the danish central graben. a maximum thickness of 106 m is found in the ravn-1 well on the southern part of the heno plateau. age. the member is of late kimmeridgian age (andsbjerg & dybkjær 2003, this volume). farsund formation history. the farsund formation was defined by vollset & doré (1984) in the norwegian north sea area and extended to the danish central graben by jensen et al. (1986). new correlations have shown that jensen et al. (1986) placed the lower boundary of the formation at a lower stratigraphic level than that defined by vollset & doré (1984). in addition, forsberg et al. (1993) pointed out a discrepancy between the base of the farsund formation as picked by vollset & doré (1984) in their type well (2/7-3) and that picked in their reference well (2/8-3); neither of these picks correspond with that suggested by jensen et al. (1986). the practice of jensen et al. (1986) is followed here as they defined the boundary at a distinct and easily recognisable log change that marks a significant change in depositional style, which is probably related to changes in subsidence patterns (andsbjerg & dybkjær 2003, this volume). however, the upper part of the formation is missing from the u-1 well that was chosen as the danish reference well for the farsund formation by jensen et al. (1986). a new reference well, the g-1 well, is thus introduced here. type and reference sections. the norwegian 2/7-3 well, 3626–3414 m b. kb, is the type section; the g-1 well in the danish central graben, 12036–7863 ft b. kb, 3631–2359 m b. msl, is designated as a reference section in the danish sector (fig. 26). a correlation line showing the stratigraphic relationship between the norwegian reference well 2/8-3 and three danish wells including the reference well g-1 and the old reference well u-1 is depicted in figure 27. lithology and environment. the farsund formation consists of medium to dark grey claystones; they are carbonaceous and variably calcareous, and are intercalated with numerous thin beds of brownish dolomite. the organic matter is mainly liptinitic, and deposition took place in a relatively deep marine environment. thin units of turbidite sandstones occur locally in the deeper parts of the basins. towards the eastern part of the danish central graben, close to the coffee soil fault, the proportion of sandstones increases and there appears to be a transition locally to the sandy poul formation. 191 192 fa rs un d vo lg ia n r ya za ni an bo m b 7863 8000 8073 8121 2423 2438 8500 9000 2359 9500 10000 k im m er id gi an vo lg ia n å sg ar d c hr on os tr at ig ra ph y fo rm at io n m b. mslft b. kb g-1 reference section: farsund formation li th ol og y 0 150 80 40 c hr on os tr at ig ra ph y fo rm at io n li th ol og y gr api sonic msec/ft m b. mslft b. kb 0 150 80 40 gr api sonic msec/ft fig. 26. the farsund formation in the g-1 well in the southern part of the danish central graben, 3631–2359 m b. msl, designated as the new reference section in the danish sector. for legend, see fig. 5. 193 fa rs un d lo la k im m er id gi an fa rs un d k im m er id gi an 10500 11000 11500 12000 12036 3631 c hr on os tr at ig ra ph y fo rm at io n m b. mslft b. kb li th ol og y 0 150 80 40 gr api sonic msec/ft c hr on os tr at ig ra ph y fo rm at io n m b. mslft b. kb li th ol og y 0 150 80 40 gr api sonic msec/ft 194 n2/8-3 baron-2 g-1 u-1 denmarknorway ? je ns en e t a l. 19 86 fa rs un d fo rm at io n lo la f m fa rs un d fm mandal fm vo lls et & d or é 19 84 h au ge su nd f or m at io n fa rs un d fo rm at io n lo la f m gr sonic gr sonic gr sonic gr sonic 200 m 50 km n 2/8-3 baron-2 g-1 u-1 log characteristics. the formation is characterised by relatively high gamma-ray values and relatively low sonic velocities. high velocity peaks and corresponding low gamma-ray readings reflecting dolomite or limestone beds or carbonate-cemented sandstone and siltstone beds are a common feature of the formation. boundaries. the farsund formation overlies the lola formation in most of the danish central graben (figs 26, 27). on plateau areas in the western and northwestern part of the danish central graben, the farsund formation overlies the heno formation (figs 21–25). the boundary with the lola formation is defined above a velocity minimum and a corresponding gamma-ray maximum. above the boundary, there is commonly a general increase in velocity, a decrease in gamma-ray values and a much higher frequency of sonic velocity peaks. the boundary with the heno formation is characterised by an abrupt or gradual increase in gammaray readings and a corresponding decrease in velocity. the upper boundary with the åsgard formation (cromer knoll group) is placed at the first significant change from high gamma-ray and low sonic velocity readings of the farsund formation to the low gammaray values and higher sonic velocities of the åsgard formation (figs 26–28). distribution and thickness. the formation extends from the norwegian north sea area, through the danish central graben. southwards it correlates with the upper part of the kimmeridge clay formation in the dutch north sea area. eastwards, the formation correlates with the børglum formation of the norwegian–danish basin, although the transition zone has not yet been located. maximum thicknesses of more than 3000 m occur in the tail end graben. age. the farsund formation is of late kimmeridgian – ryazanian age (poulsen 1986; heilmann-clausen 1987; andsbjerg & dybkjær 2003, this volume). subdivision. organic-rich mudstones in the upper levels of the farsund formation, previously referred to informally as the ‘hot unit’, are herein defined as the bo member of the farsund formation. the remainder of the formation is not subdivided at the level of members. bo member new member history. jensen et al. (1986) described the ‘hot unit’ as an informal member of the farsund formation characterised by mudstones that display unusually high gamma-ray values. correlatives at least in part are known from the norwegian sector where they are referred to the mandal formation (vollset & doré 1984), and from the dutch sector where they form the clay deep member of the kimmeridge clay formation (van adrichem boogaert & kouwe 1994). however, both the mandal formation and the clay deep member span a greater stratigraphic interval and include mudstones showing intermediate gamma-ray values beneath and above the ‘hot unit’ as described by jensen et al. (1986). use of the mandal formation in the danish sector is considered impractical due to difficulties in recognising the lower boundary in many of the danish wells. the ‘hot unit’ of jensen et al. (1986) is thus retained and given formal status as the bo member of the farsund formation. name. the bo member is named after the bo-1 well, located in the southern tail end graben where the member is well-developed. type section. the danish bo-1 well, 8561–8434 ft b. kb, 2576–2537 m b. msl (fig. 28). reference section. the danish e-1 well, 9853–9771 ft b. kb, 2966–2940 m b. msl (ineson et al. 2003, this volume, fig. 6). lithology and environment. the member is dominated by black to dark grey-brown, laminated claystones, which are carbonaceous and slightly calcareous to noncalcareous. the total organic carbon content ranges from 3–8%, locally attaining values of more than 15% (ineson et al. 2003, this volume). the organic matter is mainly liptinitic. thin beds of dolomite may occur. in the jeppe-1 core, the unit also includes thin sandstone–mudstone couplets up to 5 cm thick; log data 195 facing page: fig. 27. well-log panel from the norwegian 2/8-3 well (norwegian reference well of the farsund formation) through the baron-2 well, the g-1 well (new danish reference well of the farsund formation) and the u-1 well (previous danish reference well of the farsund formation). note the position of the basal boundary of the farsund formation as picked in the norwegian reference well and in the danish wells. suggest that thicker sandstone-rich intervals are also present in the member (ineson et al. 2003, this volume). deposition occurred in a low-energy, oxygen-deficient deep marine environment. the sandstone–mudstone couplets were deposited from dilute turbidity currents (ineson et al. 2003, this volume). log characteristics. the member is primarily identified by its high gamma-ray values, which are significantly higher than those of the underlying and overlying claystones of the farsund formation. the sonic velocity is low. the gamma-ray values may show significant variation within the member with decreasing-upwards trends, 3–5 m thick, separated by intervals with more consistently high values (e.g. bo-1). boundaries. the lower and upper boundaries are placed at shifts to the significantly lower gamma-ray values of the farsund formation. distribution and thickness. the member is recognised widely in the danish central graben, where this portion of the farsund formation is preserved. on structural highs, the upper part of the farsund formation is commonly truncated and the member may be absent or reduced in thickness. the thickness varies greatly from less than 10 m in the southern salt dome province to more than 100 m in the western part of the danish central graben, probably controlled by local factors such as structural position and sediment supply (ineson et al. 2003, this volume). age. late volgian – late ryazanian, mainly occurring within the lower ryazanian h. kochi chronozone with 196 8700 8800 8900 ft b. kb 8561 8434 8348 8400 8500 8600 li th ol og y c hr on os tr at ig ra ph y fo rm at io n m em be r 0 180 ?? r ya za ni an –v al an gi ni an 140 60 vo lg ia n fa rs un d 2576 m b. msl 2537 2511 å s. bo bo-1 type section: bo member gr api sonic msec/ft fig. 28. bo-1 well, located in the danish central graben; type section of the new bo member (farsund formation), 2576–2537 m b. msl. ås., åsgard. modified from jensen et al. (1986); for legend, see fig. 5. fa rs un d fa rs un d po ul 28189356 9400 9300 9463 9500 9600 2851 m b. mslft b. kb vo lg ia n c hr on os tr at ig ra ph y fo rm at io n gr api sonic msec/ftli th ol og y 0 140 140 40 v-1 type section: poul formation fig. 29. v-1 well, located in the eastern part of the danish central graben; type section for the poul formation, 2851–2818 m b. msl. modified from jensen et al. (1986); for legend, see fig. 5. a maximum age range from the c. preplicomphalus to the c. stenomphalus chronozones (birkelund et al. 1983; poulsen 1992; dybkjær 1998; ineson et al. 2003, this volume). poul formation history. the poul formation was defined by jensen et al. (1986). type section. the v-1 well, 9463–9356 ft b. kb, 2851–2818 m b. msl, in the eastern part of the danish central graben was selected by jensen et al. (1986; fig. 29). lithology and environment. the poul formation typically consists of dark grey siltstones and light grey, very fine-grained sandstones. the siltstones are partly argillaceous and grade into claystones. in the ugle-1 well, located at the eastern margin of the danish central graben close to the coffee soil fault zone, the formation includes pebble conglomerates, coarse-grained sandstones and fineto very fine-grained, silty sandstones. shelly debris and lithic clasts are common, and the conglomerates may be polymictic. the clasts represent a wide range of lithologies suggesting a metamorphic source terrain with phyllites and other metasediments. deposition was probably from turbidity currents on fault-bounded plateaus along the coffee soil fault. the exceptionally thick and immature succession in the ugle-1 well was probably deposited as a faultattached, localised slope apron. log characteristics. the poul formation is characterised by low gamma-ray values and intermediate sonic values representing siltand sandstones, with a few peaks showing higher sonic velocities representing carbonatecemented beds. boundaries. the formation interfingers with the farsund formation, and the lower and upper boundaries are clearly defined at abrupt changes to lower gamma-ray values and higher sonic velocities than those that characterise the farsund formation. distribution and thickness. the formation is present in the danish central graben close to the eastern faulted margin of the central graben. the thickest sandstone units occur on down-faulted plateaus along the coffee soil fault zone. the maximum thickness of 175 m is observed in the ugle-1 well (2591–2416 m b. msl). age. the poul formation is of volgian age (hoelstad 1986b; poulsen 1986). stratigraphic evolution the jurassic succession of denmark, from bornholm in the east to the central graben in the west, provides a detailed record of major basin evolutionary events that were experienced along the length of the jurassic rift system. a summary of the jurassic stratigraphic evolution of denmark is given below; detailed accounts are presented elsewhere in this volume (andsbjerg 2003, this volume; andsbjerg & dybkjær 2003, this volume; johannessen 2003, this volume; nielsen 2003, this volume). early jurassic marine deposition the early jurassic transgression of the danish basin was initiated in the norian (late triassic) and recorded by deposition of restricted marine claystones of the vinding formation and paralic – shallow marine sandstones and mudstones of the overlying and partly interfingering gassum formation. mudstones of the lower jurassic fjerritslev formation conformably overlie the gassum formation over most of the danish basin and sorgenfrei–tornquist zone, indicating a further deepening of the basin (figs 30–32). the fjerritslev formation is subdivided into four members, all of which are present in the danish basin. the two sub-units of the lower member (f-ia, f-ib) are also recognised in the danish central graben (michelsen 1978b). the formation is patchily preserved in the southern part of the norwegian north sea sector and has been drilled in the egersund subbasin and in the area around the southern vestland arch (vollset & doré 1984). to the east, comparable mudstones have been exposed at gantofta, skåne, where they are referred to the pankarp member of the rya formation (fig. 2; frandsen & surlyk 2003, this volume). the fjerritslev formation was probably deposited over most of the south-eastern north sea. the isolated occurrences of the formation are erosional outliers of a previously widely distributed lower jurassic cover, which existed prior to early middle jurassic uplift and erosion of the mid north sea dome and the ringkøbing–fyn high (michelsen 1978a, b; gowers & sæbøe 1985; ziegler 1990; underhill & partington 1993; andsbjerg & dybkjær 2003, this volume; nielsen 2003, this volume). 197 198 10 0 m m ej ru p1 h yl le bj er g1 ve ds te d1 bø rg lu m -1 fl yv bj er g1 fr ed er ik sh av n2 g r g r so ni c so ni c sp r es sp r es sp r es sp r es flv flv flll fll flb flb flll fll flb fla fla fl yv b. –h al d. fj er ri ts le v fm bø rg lu m f m fjerritslev fm m ar in e m ud st on es sh al lo w m ar in e sa nd st on es a nd s ilt st on es pa ra lic a nd n on -m ar in e sa nd st on es , si lts to ne s, m ud st on es a nd c oa ls d an is h ba si n so rg en fr ei –t or nq ui st z on e sk ag er ra k– k at te ga t pl at fo rm sk ag er ra k fm (t ri as si c) fr ed er ik sh av n fm bø rg lu m f m v ed st ed f m g as su m f m (t ri as si c) flll fll flb fla g as su m f m (ju ra ss ic ) fr ed er ik sh av n fm fl yv bj er g fm h al da ge r sa nd f m v ed st ed f m flv flll fll flb fr ed er ik sh av n2 bø rg lu m -1 fl yv bj er g1 ve ds te d1 h yl le bj er g1 m ej ru p1 50 k m fi g. 3 0. w el llo g p an el a cr o ss t h e d an is h b as in a n d f en n o sc an d ia n b o rd er z o n e sh o w in g th e u p p er t ri as si c – lo w er c re ta ce o u s lit h o st ra tig ra p h ic u n its w ith g en er al is ed d ep o si tio n al e n vi ro n m en ts i n d ic at ed b y co lo u r. f ly vb ., fl yv b je rg f o rm at io n ; h al d ., h al d ag er s an d f o rm at io n . 199 10 0 m d an is h ba si n so rg en fr ei –t or nq ui st z on e r øn ne g ra be n bø rg lu m f m flv flll fll flb fla g as su m f m (t ri as si c) fjerritslev fm flll fll flb fla fj er ri ts le v fm h yl le bj er g1 te rn e1 h an s1 pe rn ill e1 st in a1 g as su m f m (t ri as si c) fr ed er ik sh av n fm q ua te rn ar y g as su m f m (ju ra ss ic ) g as su m f m (t ri as si c) u pp er c re ta ce ou s h as le f m rønne fm rønne fm m un ke ru p m b m un ke ru p m b q ua te rn ar y so rt ha t fm h as le f m tr ia ss ic tr ia ss ic fl yv bj er g fm h al da ge r sa nd f m so se bu gt m b g al ge lø kk e m b g al ge lø kk e m b so se bu gt m b ve ds te d fm m ar in e m ud st on es sh al lo w m ar in e sa nd st on es a nd s ilt st on es pa ra lic a nd n on -m ar in e sa nd st on es , si lts to ne s, m ud st on es a nd c oa ls h yl le bj er g1 te rn e1 h an s1 pe rn ill e1 st in a1 g r so ni c g r so ni c so ni c g r g r so ni c g r so ni c 50 k m fi g. 3 1. w el llo g p an el f ro m t h e d an is h b as in a n d f en n o sc an d ia n b o rd er z o n e w ith t h e u p p er t ri as si c – lo w er c re ta ce o u s lit h o st ra tig ra p h ic u n its c o m p ar ed t o t h e p er n ill e1 an d s tin a1 w el ls fr o m t h e r ø n n e g ra b en , o ff sh o re b o rn h o lm . g en er al is ed d ep o si tio n al e n vi ro n m en ts i n d ic at ed b y co lo u r. the early jurassic transgression resulted in the establishment of a well-oxygenated marine environment in the danish basin and in the central north sea basin (fig. 33; michelsen 1975, 1978a; pedersen 1986). marine claystones interbedded with thin sandstone beds and carbonate-cemented beds (f-ia unit) are known from both areas. along the north-eastern margin of the danish basin, fluvial and shallow marine conditions prevailed in the hettangian – early sinemurian, and a stepwise backstepping of the sandy gassum formation took place (figs 30–32; nielsen 2003, this volume). in skåne, southern sweden, the hälsingborg and döshult members represent a transition from deltaic to shallow marine conditions (sivhed 1984; erlström et al. 1999; ahlberg et al. 2003, this volume). on bornholm, an overall transgressive succession of lacustrine, coastal plain and tidal flat deposits is represented by the rønne formation, which is thickly preserved in the rønne graben (onshore: galgeløkke-1, -2; offshore: pernille-1, stina-1; figs 5, 31; gravesen et al. 1982; nielsen 1995; surlyk et al. 1995). the depositional environment changed in the early sinemurian due to continued sea-level rise. deposition of fine-grained claystones (f-ib unit) took place from the early sinemurian to the earliest early pliensbachian in the basinal areas. to the north, in jylland, the gassum formation was overstepped, and deposition of marine clay (f-ib unit) extended onto the skagerrak–kattegat platform (fig. 30). deposition of fluvial and shallow marine sandy deposits of the döshult formation in skåne was followed by deposition of marine clays of the pankarp member, whereas shallow marine sand, tidal flat heteroliths and estuarine mud are found in the correlative part of the rønne formation on bornholm (surlyk et al. 1995; erlström et al. 1999; frandsen & surlyk 2003, this volume; nielsen 2003, this volume). continued sea-level rise in the early pliensbachian led to open marine conditions and deposition of silty clay (f-ii member) in the danish basin and more sandy deposits on the skagerrak–kattegat platform (fig. 30). farther to the east and south-east, in the fennoscandian border zone, marine sand and clay (katlösa member in skåne) and marine sand and silt (hasle formation on bornholm) were deposited (troedsson 1951; surlyk & noe-nygaard 1986; norling et al. 1993; frandsen & surlyk 2003, this volume; nielsen 2003, this volume). after a regressive event with deposition of minor sands and erosion on the skagerrak–kattegat platform, deposition of clay again dominated in the danish basin (f-iii member) in late pliensbachian – early toarcian times (fig. 30). at the end of the pliensbachian, the benthic faunas disappeared due to increasing anoxia, probably related to a sea-level rise, corroborated by sedimentology, organic geochemical analyses and palynology (pedersen 1986; michelsen 1989b; dybkjær 1991; nielsen 2003, this volume). the sea-level rise culminated in the early toarcian. towards the end of the toarcian, a regression began with deposition of interbedded clay and sand (f-iv member) in a more shallow and restricted marine environment (figs 30, 31). a regressive tendency also occurred further south in the fennoscandian border zone during the late pliensbachian as reflected by dark grey, grey-green and red-brown marine sandstones and siltstones (rydebäck member) in skåne. the benthic foraminiferal fauna decreased markedly in density and diversity from the pliensbachian to the toarcian, but ammonites occur throughout the formation (norling 1972). on bornholm, sandy and clayey deposits interbedded with thin coal beds (the sorthat formation) were deposited in brackish coastal plain environments (koppelhus & nielsen 1994; petersen & nielsen 1995; petersen et al. 2003b, this volume). the lower–middle jurassic boundary is difficult to identify in the danish basin due to relatively poor biostratigraphic data and is conventionally placed between the fjerritslev and haldager sand formations (michelsen 1978a, 1989a). recent studies indicate that the lower– middle jurassic transition is situated within the uppermost part of the fjerritslev formation in the sorgenfrei–tornquist zone; elsewhere, it coincides with the marked erosion surface between the fjerritslev and haldager sand formations (nielsen 2003, this volume). the maximum thicknesses of the uppermost member (f-iv) of the lower jurassic fjerritslev formation and the middle jurassic haldager sand formation are found in the sorgenfrei–tornquist zone, indicating continuous subsidence contemporaneously with the middle jurassic uplift of the ringkøbing–fyn high (michelsen & nielsen 1991; nielsen 1992, 1994, 2003, this volume). the gradual eastwards transgression during the early jurassic known from the fennoscandian border zone is also recorded in north-eastern germany (meinhold et al. 1960; bertelsen 1978; michelsen 1978a). to the south-west, in the dutch central graben, a continuous succession of lower–middle jurassic deposits is included in the altena group, comprising the aalburg, posidonia shale and werkendam formations (fig. 34; nam & rgd 1980; van adrichem boogaert & kouwe 1994; herngreen et al. 2003, this volume). the overall lithology of the altena group is similar to that of the fjerritslev formation. the aalburg formation is of hettangian– pliensbachian age, and the sediments seem to have 200 201 c le o1 l1 in ez -1 f1 k -1 bø rg lu m -1 g r u . j ur . tr ia ss ic g as su m f m h al da ge r sa nd fm bø rg lu m f m g as su m f m t ri as si c f lv v ed st ed f m fr ed er ik sh av n fm bø rg lu m f m fl yv bj er g fm h al da ge r sa nd f m flv flll fll flb flb fla gassum fm (jurassic)fjerritslev fm fe lic ia -1 j1 10 0 m c en tr al g ra be n r in gk øb in g– fy n h ig h n or w eg ia n– d an is h ba si n so rg en fr ei –t or nq ui st z on e so ni c so ni c g r g r so ni c so ni c g r sp r es g r so ni c so ni c g r g r so ni c m ar in e m ud st on es tu rb id ite s an ds to ne s an d si lts to ne s pa ra lic a nd n on -m ar in e sa nd st on es , si lts to ne s, m ud st on es a nd c oa ls sh al lo w m ar in e sa nd st on es a nd s ilt st on es bø rg lu m -1 fe lic ia -1 j-1 k -1 f1 in ez -1 c le o1 l1 fr ed er ik sh av n fm lo w er c re t. fl yv bj er g fm c ha lk g r br yn e fm lu lu fm t ri as si c lo la fm fa rs un d fm 50 k m fi g. 3 2. w el llo g p an el r u n n in g fr o m t h e so rg en fr ei – to rn q u is t z o n e to t h e n o rt h -e as te rn d an is h c en tr al g ra b en s h o w in g th e u p p er t ri as si c – lo w er c re ta ce o u s lit h o st ra tig ra p h ic u n its w ith g en er al is ed d ep o si tio n al en vi ro n m en ts i n d ic at ed b y co lo u r. 202 4° e 8° e 12 °e 16 °e 58 °n 5 6° n la te a al en ia n – ba th on ia n 4° e 8° e 12 °e 16 °e 58 °n 56 °n h et ta ng ia n 4° e 8° e 12 °e 16 °e 58 °n 56 °n si ne m ur ia n – ea rl y pl ie ns ba ch ia n 4° e 8° e 12 °e 16 °e 58 °n 56 °n to ar ci an – ea rl y a al en ia n sw ed en n or w ay g er m an y t he n et he rl an ds sw ed en n or w ay g er m an y t he n et he rl an ds sw ed en n or w ay g er m an y t he n et he rl an ds sw ed en n or w ay g er m an y d en m ar k d en m ar k t he n et he rl an ds 10 0 km 10 0 km 10 0 km 10 0 km d en m ar k d en m ar k 203 d en m ar k o ffs ho re m ar in e (m ud -d om in at ed ) d ee p m ar in e (s an ds , g ra ve ls ) n on -d ep os iti on /e ro si on pa ra lic a nd n on -m ar in e sa nd st on es , si lts to ne s, m ud st on es a nd c oa ls sh al lo w m ar in e sa nd st on es a nd s ilt st on es 4° e 8° e 12 °e 16 °e 58 °n 56 °n c al lo vi an 4° e 8° e 12 °e 16 °e 58 °n 56 °n k im m er id gi an 4° e 8° e 12 °e 16 °e 58 °n 56 °n vo lg ia n – ea rl y r ya za ni an 10 0 km sw ed en n or w ay g er m an y t he n et he rl an ds sw ed en n or w ay g er m an y t he n et he rl an ds sw ed en n or w ay g er m an y t he n et he rl an ds 10 0 km 10 0 km d en m ar k d en m ar k fi g. 3 3. p al ae o ge o gr ap h ic m ap s sh o w in g th e in fe rr ed d is tr ib u tio n o f gr o ss d ep o si tio n al e n vi ro n m en ts d u ri n g th e ju ra ss ic . been deposited in a lower shoreface to deep offshore environment similar to that of the lower fjerritslev formation. the posidonia shale formation is of early toarcian age and seems to have been deposited under dysoxic conditions as were the upper part of the f-iii and the lower part of the f-iv members of the fjerritslev formation. the werkendam formation was deposited in a shallow, open marine environment during late toarcian – earliest bathonian times. middle jurassic uplift during late aalenian – early bajocian times, the central north sea area including the ringkøbing–fyn high became uplifted and formed a broad arch (michelsen 1978a; koch 1983; ziegler 1990; andsbjerg et al. 2001; nielsen 2003, this volume) rather than a dome as suggested by underhill & partington (1993, 1994). the uplift influenced a large area from england to denmark and widespread erosion occurred. the sorgenfrei– tornquist zone seems to have been mainly unaffected by uplift and erosion, but farther south and south-westwards in the danish basin, erosion cut deeply into the fjerritslev formation and totally removed these strata from the ringkøbing–fyn high (michelsen 1978a; koch 1983; nielsen 2003, this volume). on the eastern part of the ringkøbing–fyn high, erosion also removed the upper part of the triassic succession, and triassic sediments are absent from the western part of the high. the main rifting of the central graben was initiated during the middle jurassic subsequent to the uplift of the central north sea (fig. 33). down-faulting took place along the eastern bounding coffee soil fault leading to preservation of lower jurassic deposits in the south-eastern part of the danish central graben. erosion cut deeply into the pre-jurassic deposits in the north-western part of the central graben due to simultaneous half-graben tilting, and permian and carboniferous rocks subcrop postlower jurassic deposits in this area (fig. 35). middle jurassic block-faulting and subsidence characterised the sorgenfrei–tornquist zone, where the entire lower jurassic – lowermost middle jurassic succession is preserved (michelsen & nielsen 1991; nielsen 2003, this volume). the regional uplift also affected skåne where renewed faulting and local uplift accompanied by volcanism clearly influenced sedimentation in middle jurassic time by reducing the depositional area to a narrow basin (klingspor 1976; rolle et al. 1979; norling & bergström 1987; norling et al. 1993; ahlberg et al. 2003, this volume). middle jurassic fluvial and paralic deposition the early middle jurassic regional uplift of the central north sea basin resulted in radical palaeogeographic changes, including closure of the seaway linking the arctic sea with the tethys ocean. in the danish area, subsequent rifting changed the basin configuration from the extensive early jurassic sea into two fault-controlled zones of deposition, the central graben and the danish basin where a depocentre was situated in the sorgenfrei– tornquist zone (fig. 33; andsbjerg et al. 2001). regional and local erosion cut deep into the lower jurassic and older deposits. the erosional surface was onlapped by fluvial, paralic and marine sediments during middle–late jurassic times owing to renewed regional subsidence (andsbjerg & dybkjær 2003, this volume; johannessen 2003, this volume; nielsen 2003, this volume). middle jurassic deposits occur in the central and north-eastern parts of the danish basin. there is a marked unconformity between the marine claystones of the lower jurassic fjerritslev formation and the fluvial to shallow marine sandstones of the middle jurassic haldager sand formation. in the fault-bounded sorgenfrei–tornquist zone, the boundary between the formations occurs in the lower aalenian and is developed as a marine regressive erosion surface (poulsen 1996; nielsen 2003, this volume). the main change in palaeogeography involving uplift, regression and influx of sand thus seems to have occurred in the early aalenian. subsequent to the uplift event, a middle jurassic succession, more than 150 m thick, accumulated in the sorgenfrei–tornquist zone; this succession comprises fineto coarse-grained, shallow marine sandstones and fluvial sandstones and siltstones interbedded with marine and lagoonal claystones and thin coal beds (haldager sand formation, figs 30–32). this succession thins to 204 facing page: fig. 34. comparison of the upper triassic – lower cretaceous lithostratigraphic schemes from the north sea central graben, the norwegian–danish basin and the fennoscandian border zone. sst., sandstone; m.w.mb, middle werkendam member; m.g.mb, middle graben sandstone member; u.g., upper graben formation. modified from bertelsen (1978, 1980); michelsen (1978, 1989a); nam & rgd (1980); gravesen et al. (1982); sivhed (1984); vollset & doré (1984); herngreen & wong (1989); norling et al. (1993); richards et al. (1993); van adrichem boogaert & kouwe (1994); koppelhus & nielsen (1994); ahlberg et al. (2003, this volume); herngreen et al. (2003, this volume). 205 triassicjurassiccretaceous upper upperlower lowermiddleseries system st ag e upper upperlower lowermiddleseries system triassicjurassiccretaceous n or th s ea c en tr al g ra be n n s w br iti sh se ct or n or w eg ia n se ct or d an is h se ct or d ut ch se ct or fi sk eb an k su bb as in d an is h ba si n sk ån e bo rn ho lm n or w eg ia n– d an is h ba si n fe nn os ca nd ia n bo rd er z on e n e s n n n s s s st ag e v al an gi ni an v al an gi ni an r ya za ni an r ya za ni an v ol gi an v ol gi an k im m er id gi an k im m er id gi an o xf or di an o xf or di an c al lo vi an c al lo vi an ba th on ia n ba th on ia n ba jo ci an ba jo ci an a al en ia n a al en ia n t oa rc ia n t oa rc ia n pl ie ns ba ch ia n pl ie ns ba ch ia n si ne m ur ia n si ne m ur ia n h et ta ng ia n h et ta ng ia n r ha et ia n r ha et ia n n or ia n n or ia n v lie la nd c la ys to ne fo rm at io n c la y d ee p m em be r k åg er öd f or m at io n r yd eb äc k m em be r k at sl ös a m em be r pa nk ar p m b d ös hu lt m b h el si ng bo rg m em be r bj uv m b va llå kr a m b höganäs fmrya fm g la ss s an d m b fu gl un da m b fo rt un a m ar l fy le da le n c la y annero formation v ita bä ck c la y n yt or p sa nd jy de gå rd fo rm at io n r ob be da le fo rm at io n r ab ek ke fo rm at io n ba gå fo rm at io n so rt ha t fo rm at io n h as le f or m at io n rønne formation g al ge lø kk e m b so se b ug t m b m un ke ru p m b r is eb æ k m b kågeröd fm fj er ri ts le v fo rm at io n w in te rt on fo rm at io n va lh al l fo rm at io n ve ds te d fo rm at io n fr ed er ik sh av n fo rm at io n bø rg lu m fo rm at io n fl yv bj er g fo rm at io n h al da ge r sa nd fo rm at io n fj er ri ts le v fo rm at io n g as su m fo rm at io n v in di ng fm sk ag er ra k fm o dd es un d fm va lh al l f m å sg ar d fm fa rs un d fm u la fm m an da l fo rm at io n k im m er id ge c la y fo rm at io n r ib bl e ss t.m b h ea th er fm fr es hn ey ss t.m b fulmar fm fulmar fm ? ? rattray volcanic member ron volcanic member pentland fm fjerritslev formation fjerritslev formation ?? ? ? ? ? ? ? ? ? ? sk ag er ra k fo rm at io n sm ith b an k fo rm at io n g as su m fo rm at io n br yn e fo rm at io n h au ge su nd fo rm at io n el df is k fm å sg ar d fm le ek m b v yl fm bo m b fa rs un d fo rm at io n po ul fm lu lu f m m id dl e g ra be n fm h en o fm lo la fo rm at io n br yn e fo rm at io n sc ru ff g re en sa nd f m k im m er id ge c la y fo rm at io n pu zz le h ol e fm u .g m id dl e g ra be n fm m .g . m b lo w er g ra be n fm w er ke nd am fm w er ke nd am fmm . w . m b po si do ni a sh al e fo rm at io n a al bu rg fo rm at io n sl ee n fo rm at io n u pp er k eu pe r c la y m b d ol om iti c k eu pe r m em be r r ed k eu pe r c la y m b sm ith b an k m em be r sk ag er ra k fo rm at io n g as su m fo rm at io n fj er ri ts le v fo rm at io n br yn e fo rm at io n sa nd ne s fo rm at io n eg er su nd fo rm at io n ta u fo rm at io n sa ud a fo rm at io n fl ek ke nf jo rd fo rm at io n ? ? mariedal fm 206 10 0 m d ee p g or m -1 el ly -2 r av n1 g w en -2 je pp e1 g er t4 1 c ro m er k no ll g p fa rs un d fm fa rs un d fm h en o fm h en o fm lo la f m tr ia ss ic lo la f m fa rs un d fm br yn e fo rm at io n h en o fm lo la f m fj er ri ts le v fm m id dl e g ra be n fm tr ia ss ic c ro m er k no ll g p z ec hs te in g p m ar in e m ud st on es m ar in e or ga ni cri ch m ud st on es tu rb id ite s an ds to ne s an d si lts to ne s sh al lo w m ar in e sa nd st on es a nd s ilt st on es pa ra lic a nd n on -m ar in e sa nd st on es , s ilt st on es , m ud st on es a nd c oa ls 50 k m d ee p g or m -1 el ly -2 r av n1 g w en -1 je pp e1 g er t4 fe da g ra be n h en o pl at ea u sa lt d om e pr ov in ce g r so ni c r es g r so ni c r es g r so ni c r es g r so ni c r es g r so ni c r es g r so ni c r es less than 10 m in central and south-western parts of the danish basin. on the skagerrak–kattegat platform, north-east of the sorgenfrei–tornquist zone, the thickness of the haldager sand formation varies between 40 m and 90 m. a braided river system running downslope from the uplifted ringkøbing–fyn high has been suggested as the main method of transport for the sediments (koch 1983). the middle jurassic deposits are generally poor in fossils, but palynological assemblages suggest an aalenian–callovian age (reviewed by nielsen 2003, this volume). dinoflagellate cysts from the top part of the haldager sand formation in the vedsted-1 well section indicate a late callovian or younger age (poulsen 1992, 1996), suggesting that the youngest sediments of the haldager sand formation in the sorgenfrei–tornquist zone were deposited during a callovian sea-level rise. in the north-western part of the norwegian–danish basin and in the central graben, sediments equivalent to the haldager sand formation are referred to the fluvio-deltaic bryne formation. to the south-east, along the fennoscandian border zone in skåne, southern sweden, the development corresponds to that of the sorgenfrei–tornquist zone in the danish area. a change from the marine rydebäck formation to fluvial, lacustrine, lagoonal and shallow marine deposits of the fuglunda beds (mariedal formation) and the vilhelmsfält formation (ahlberg et al. 2003, this volume) took place in the late aalenian, and deltaic deposition seems to have continued throughout the bathonian (figs 2, 33, 34). the aalenian part of the rydebäck formation consists of black and variegated siltstones with a varying clay content showing a general regressive tendency, and the foraminiferal fauna is sparse compared to that of the underlying lower jurassic succession (norling 1972). several conglomeratic horizons suggest breaks in marine deposition (norling et al. 1993). the fuglunda beds consist of alternating sand and clay with thin coal beds. deposition of clean sand, probably of shallow marine origin (the glass sand member of the mariedal formation), took place during most of the bathonian. the clay-dominated, marine fortuna marl was deposited in the southern part of skåne from latest bathonian time, throughout the callovian and into the early oxfordian (norling 1972). marine influence in the middle jurassic thus seems to have increased south-eastwards along the fennoscandian border zone (fig. 2). on bornholm, an erosional unconformity between the sorthat and bagå formations separates lower jurassic paralic deposits from middle jurassic fluvial and lacustrine deposits (koppelhus & nielsen 1994). the middle jurassic deposits consist of fluvial sand and gravel and lacustrine clay with coal beds deposited during late aalenian(?) and bajocian–bathonian times. locally, the succession contains debris flow deposits indicating movements on the eastern bounding fault of the rønne graben in bathonian time (gry 1969). the top of the bagå formation is an erosional unconformity that forms the present-day land surface, and callovian and upper jurassic strata are absent on bornholm, although the lowermost deposits of the rabekke formation in the nyker fault block may extend down into the volgian (gravesen et al. 1982). in the danish central graben, middle jurassic deposition took place in the down-faulted area west of the coffee soil fault zone. sandstones with subordinate siltstones and claystones, and laterally extensive coal beds in the upper part of the succession, characterise the middle jurassic bryne and lulu formations in the northern danish central graben (fig. 36). an upwards increase in marine influence is reflected by a change from fluvial or coastal plain deposits in the lower part to marginal and shallow marine deposits in the upper part of the succession (johannessen & andsbjerg 1993; andsbjerg 2003, this volume). in the southern danish central graben, the lower part of the middle jurassic succession is characterised by thick sandy deposits interbedded with siltand claystones and occasional thin coal beds (bryne formation). the upper part contains claystones, siltstones and common coal beds (middle graben formation; figs 35, 36). the depositional environment evolved from a coastal plain environment with strong fluvial influence to a lowenergy paralic environment dominated by lagoons or interdistributary bays. the environmental evolution through middle jurassic time shows an increasing marine influence culminating in the major late jurassic marine transgression, which 207 facing page: fig. 35. nw–se well-log panel through the danish central graben showing the upper triassic – lower cretaceous lithostratigraphic units with generalised depositional environments indicated by colour. note the organic-rich mudstones, locally associated with turbidite sandstones, in the uppermost farsund formation of these wells (excluding gert-4); these are referred to the bo member (see above and ineson et al. 2003, this volume). lower cretaceous strata of the cromer knoll group cap the farsund formation in this transect with the exception of the elly-2 well in which the farsund formation is unconformably overlain by the upper cretaceous chalk group. led to deposition of deep offshore clay-dominated deposits. the facies and age distribution of the middle jurassic deposits recognised in central graben well sections from the norwegian, danish and dutch sectors indicate that the transgression came from the north (michelsen et al. 1987; johannessen & andsbjerg 1993; underhill & partington 1994). in the norwegian fiskebank subbasin and the norwegian sector of the central graben, the non-marine bryne formation is overlain by shallow marine and coastal deposits of the callovian sandnes formation (fig. 34; vollset & doré 1984). similarly, the first fully marine middle jurassic deposits in the danish central graben belong to the callovian lulu formation, whereas the first comparable fully marine deposits in the dutch central graben (referred to the upper graben formation) are of middle–late oxfordian age (see herngreen et al. 2003, this volume). thus the relatively sparse biostratigraphic data indicate a southwards younging of the transgression. the basal part of the norwegian haugesund and the danish lola formations are of latest callovian – early oxfordian age (vollset & doré 1984; jensen et al. 1986; andsbjerg & dybkjær 2003, this volume). the onset of marine middle to outer shelf conditions thus started in latest middle jurassic time in the central north sea basin, but only reached the southern part of the basin in the oxfordian (fig. 34; herngreen & wong 1989). late jurassic marine deepening the structural reorganisation established during early middle jurassic uplift and subsequent erosion created the framework for the basin configuration that prevailed from the middle and late jurassic into early cretaceous times. the ringkøbing–fyn high including adjacent areas to the south and north remained an uplifted landmass, forming the southern border of the 208 50 m west lulu-3 middle gr. fm fjerritslev fm bryne fm lola fm sonicgr amalie-1 elly-3 u-1 sonic sonicgr gr sonicgr lola fm lulu fm bryne fm triassic søgne basin tail end graben heno plateau salt dome province marine mudstones shallow marine sandstones and siltstones paralic sandstones, siltstones, mudstones and coals fluvial and estuarine channel sandstones and conglomerates floodplain and lacustrine mudstones, siltstones and coals 50 km u-1 elly-3 amalie-1 west lulu-3 fig. 36. north to south well-log panel through the danish central graben showing the middle – lower upper jurassic lithostratigraphic units with generalised depositional environments indicated by colour. danish basin. the landmass was submerged late in the early cretaceous. the fennoscandian border zone and the easternmost part of the basin acted as a broad transition zone characterised by shallow marine to paralic or non-marine deposition during repeated transgressive–regressive cycles. these environmental conditions lasted into the earliest early cretaceous, when they were succeeded by deeper marine conditions. the initial shallow marine conditions represented by the uppermost, probably callovian part of the haldager sand formation were succeeded by transgressive, paralic to marine deposition (flyvbjerg formation) in the oxfordian. dinoflagellate cysts from the lowermost part of the flyvbjerg formation in wells located in the sorgenfrei–tornquist zone and in a well in the central part of the danish basin suggest an early–middle oxfordian age and a marine origin for the basal part of the formation (poulsen 1992, 1996). the transition from the shallow marine, regressive deposits of the uppermost part of the flyvbjerg formation to the transgressive, deeper marine deposits of the børglum formation seems to be of earliest kimmeridgian age in the danish basin and of late kimmeridgian age in the fennoscandian border zone, thus demonstrating a north-eastwards younging (poulsen 1992, 1996) and recording the gradual transgression of the basin margin. the deepening represented by the børglum formation lasted from the earliest kimmeridgian to the middle volgian in the danish basin. a regressive development followed with deposition of coarser sediments in most of the basin. in the fennoscandian border zone, two–three coarsening-upwards successions with shallow marine silt and sand beds alternating with transgressive clays were deposited during the volgian–ryazanian (frederikshavn formation; figs 30–32). farther to the north-west, in the norwegian– danish basin, a nearly time-equivalent succession of sedimentary rocks is referred, in stratigraphic order, to the egersund, tau and sauda formations (vollset & doré 1984). to the south-east, in skåne, the lacustrine and lagoonal variegated claystones and siltstones of the fyledal clay were deposited during the oxfordian and early kimmeridgian, succeeded by prograding coastal plain deposits of the nytorp sand (norling et al. 1993; ahlberg et al. 2003, this volume). the jurassic–cretaceous transitional beds of the vitabäck clays are again dominated by brackish to freshwater deposits showing some marine influence. the connection between the central european basins and the danish basin along the fennoscandian border zone was characterised by increased marine influence (fortuna marl) during the callovian (fig. 34). during late jurassic time, marine communication between the danish basin and the north german basin was mainly through the central graben, but a connection was periodically established through the sorgenfrei–tornquist zone (christensen & kilenyi 1970). the environment changed to mainly brackish water, while open marine conditions dominated in the danish basin. however, repeated regressive–transgressive events are recorded in skåne and in northernmost jylland. precise stratigraphic correlation between these two areas is not yet possible due to scarce biostratigraphic data. callovian– kimmeridgian deposits are absent on bornholm and in the rønne graben. the ?volgian– valanginian rabekke, robbedale and jydegård formations (fig. 2), preserved east of the rønne graben, show a broadly transgressive development, probably reflecting the overall early cretaceous transgression in northwest europe. late jurassic graben development and sedimentation onset of the main rifting following the middle jurassic uplift of the central north sea basin led to rapid subsidence of the central graben. fault activity along the coffee soil fault continued from middle jurassic through late jurassic times. the central graben was developed as a major half-graben, which became compartmentalised into a number of minor grabens and highs during the late jurassic (møller 1986; japsen et al. 2003, this volume; møller & rasmussen 2003, this volume). up to 4000 m of clay-dominated sediments were deposited in the tail end graben to the east during the late jurassic (møller 1986; michelsen et al. 1992). the marine middle–outer shelf conditions which were established in latest callovian time in the danish central graben continued during the late jurassic overall sea-level rise. the central graben region was probably connected with the danish basin during most of the late jurassic through the north-western part of the norwegian–danish basin, north of the ringkøbing–fyn high. the sedimentary facies of the central graben differ markedly from those of the norwegian–danish basin, and none of the formations known from the danish basin can be traced to the central graben. these differences probably resulted from varying distances to the source areas and from differences in structural evolution. deposition of offshore muds of the lola formation in the eastern part of the central graben continued 209 210 fa rs un d fm lo la fmh en o fm tr ia ss ic 10 0 m fa rs un d fmpo ul fm c al ed on ia n ba se m en t c ro m er k no ll g p c ro m er k no ll g p tr ia ss ic lo la fo rm at io n h en o pl at ea u ta il en d g ra be n el in -1 v1 r av n1 u gl e1 g r so ni c r es g r so ni c r es g r so ni c r es g r so ni c r es m ar in e m ud st on es tu rb id ite s an ds to ne s, si lts to ne s an d co ng lo m er at es sh al lo w m ar in e sa nd st on es a nd s ilt st on es 50 k m el in -1 v1 u gl e1 r av n1 from the latest callovian, and onlapped westwards onto the heno plateau during the oxfordian–kimmeridgian (fig. 35; michelsen et al. 1992; johannessen & andsbjerg 1993; andsbjerg & dybkjær 2003, this volume). thin shallow marine sands were deposited at the western margin of the basin during deposition of the lola formation. this indicates that the sea transgressed westwards and up-dip from the eastern part of the danish central graben. deposition of offshore muds of the lola formation on the heno plateau was succeeded by shoreface sands of the heno formation during the latest early – late kimmeridgian (figs 35, 37; andsbjerg & dybkjær 2003, this volume; johannessen 2003, this volume). shallow marine and back-barrier sediments of the heno formation were also deposited on the gertrud plateau and in the southern feda graben. rapid subsidence of the individual basins of the central graben took place during latest kimmeridgian time with an increased rate in the early volgian. deposition of offshore muds dominated the entire graben area during latest kimmeridgian – volgian times, periodically with a significant increase in the content of organic matter as seen in the lower part of the farsund formation in the feda graben and heno plateau (fig. 35; damtoft et al. 1992). locally, coarse-grained sediments deposited from sediment gravity flows (poul formation) interfinger with the claystones of the farsund formation adjacent to active fault zones (fig. 37). the subsidence rates seem to have decreased in the latter part of the volgian, and during the late volgian – early ryazanian, the highly radioactive bo member developed as a part of the farsund formation. deposition of the organic-rich muds was governed by stratification of the water column, poor circulation, sea-level changes and reduced sediment influx (ineson et al. 2003, this volume). the norwegian sector of the central graben developed in a similar way to that described for the danish part. the callovian – early volgian clay deposits are included in the haugesund formation (fig. 34), which differs from the partly time-equivalent lola formation in containing a number of thin sandstone beds and by forming an overall coarsening-upwards cycle (vollset & doré 1984; see discussion in forsberg et al. 1993). the overlying farsund formation is widely distributed north of the danish sector. in the western part of the graben, a sand-dominated turbidite succession of kimmeridgian age (eldfisk formation) forms a wedge between the haugesund and farsund formations. in the norwegian sector, the farsund formation is also capped by an organic-rich mudstone, named the mandal formation (fig. 34). to the south, in the dutch sector, the callovian – upper oxfordian lower, middle and upper graben formations (terrestrial–paralic) of the dutch north sea area are succeeded, with some overlap, by the upper oxfordian – kimmeridgian puzzle hole formation (lower delta plain deposits), indicating a northwards shift of the coastline during the middle oxfordian (fig. 34; nam & rgd 1980; herngreen & wong 1989; van adrichem boogaert & kouwe 1994; herngreen et al. 2003, this volume). transgression was again initiated in the late oxfordian, as indicated by the southwards onlap by the deeper marine kimmeridge clay formation (equivalent to the lola and farsund formations) onto the shallow marine deposits. to the south, a barrier sand complex (the volgian–ryazanian scruff greensand formation) seems to have been partly reworked into transgressive sheet sands overstepping the graben margins (herngreen & wong 1989). at the jurassic– cretaceous transition, anoxic conditions (clay deep member) prevailed in the northern part of the dutch north sea area. acknowledgements the referees g.f. waldemar herngreen and gunver k. pedersen and the editor jon r. ineson are thanked for many valuable suggestions. f.s. acknowledges the support of the danish natural science research foundation; l.h.n. and f.s. acknowledge the use of gravgærde, the field station of the university of copenhagen on the island of bornholm. 211 facing page: fig. 37. west–east well-log panel from the danish central graben showing middle/upper jurassic – lower cretaceous lithostratigraphic units with generalised depositional environments indicated by colour. note the thin development of organic-rich mudstones (bo member) in the uppermost farsund formation of the ravn-1 well; this member is not recognised in the v-1 and ugle-1 wells close to the eastern flank of the central graben and is thick but only weakly developed in the elin-1 well (see ineson et al. 2003, this volume). 212 references ahlberg, a., sivhed, u. & erlström, m. 2003: the jurassic of skåne, southern sweden. 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(eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 611–630 (this volume). petersen, h.i., nielsen, l.h., koppelhus, e.b. & sørensen, h.s. 2003: early and middle jurassic mires of bornholm and the fennoscandian border zone: a comparison of depositional environments and vegetation. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 631–656 (this volume). poulsen, n.e. 1986: callovian–volgian dinocyst stratigraphy of the central trough in the danish north sea area. bulletin of the geological society of denmark 35, 1–10. poulsen, n.e. 1991: upper jurassic dinocyst stratigraphy in the danish central trough. in: michelsen, o. & frandsen, n. (eds): the jurassic in the southern central trough. danmarks geologiske undersøgelse serie b 16, 7–15. poulsen, n.e. 1992: jurassic dinoflagellate cyst biostratigraphy of the danish subbasin in relation to sequences in england and poland; a preliminary review. review of palaeobotany and palynology 75, 33–52. poulsen, n.e. 1994: dinoflagellate cyst biostratigraphy of rhaetian–ryazanian (uppermost triassic – lowermost cretaceous) deposits from the danish subbasin. geobios 17, 409–414. 216 poulsen, n.e. 1996: dinoflagellate cysts from marine jurassic deposits of denmark and poland. american association of stratigraphic palynologists contributions series 31, 227 pp. rasmussen, l.b. 1974: some geological results from the first five danish exploration wells in the north sea. danmarks geologiske undersøgelse iii. række 42, 47 pp. rasmussen, l.b. 1978: geological aspects of the danish north sea sector; with a report on the wells dansk nordsø e-1, e-2, f-1, g-1, h-1, i-1, j-1, and k-1. danmarks geologiske undersøgelse iii. række 44, 85 pp. richards, p.c., lott, g.k., johnson, h., knox, r.w.o’b. & riding, j.b. 1993: jurassic of the central and northern north sea. in: knox, r.w.o’b. & cordey, w.g. 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(eds) 1984: a revised triassic and jurassic lithostratigraphic nomenclature for the norwegian north sea. norwegian petroleum directorate bulletin 3, 53 pp. ziegler, p.a. 1975: geologic evolution of the north sea and its tectonic framework. american association of petroleum geologists bulletin 59, 1073–1097. ziegler, p.a. 1990: geological atlas of western and central europe, 2nd edition, 239 pp. amsterdam: elsevier for shell internationale petroleum maatschappij. manuscript received 7 march 2000; revision accepted 25 june 2001. geological survey of denmark and greenland bulletin 23, 2011, 17–2 17 the submarine siri canyon is ne–sw-oriented and located in the danish north sea (fig. 1). it contains a number of oil reservoirs with glauconite-rich sand. the reservoirs of interest in the nini oil field are the late paleocene tyr member of the lista formation and the kolga member of the sele formation (schiøler et al. 2007), presumably of early eocene age. these members have previously been known as the ty and hermod members (hamberg et al. 2005; poulsen et al. 2007). the sand shows signs of injection, both in cores and in seismic data. the aim of this work is to chemically characterise and fingerprint the sand in order to reveal the origin of the sand found in three horizontal wells, which could have been injected from one or both of the tyr and kolga members. core samples were collected from two vertical wells of known stratigraphy to make a basis of comparison, whereas samples of the cuttings were collected from the three horizontal wells with ages primarily corresponding to the kolga member. the purpose was moreover to evaluate whether cuttings samples can be used for fingerprinting as an alternative to core samples. the interest in discriminating between the ages of the injected sand is the fact that the reservoir properties (porosity and permeability) are largely controlled by the original composition of the sand. consequently, results from this study could affect the property modelling of the field. sand from the tyr and kolga members is dominated by quartz and glauconite and contains fairly well-preserved kfeldspar, plagioclase and mica. the content of feldspar and mica is quite constant, and the feldspar and quartz grains are equally rounded. k-feldspar is more common and better preserved than plagioclase, and k-feldspar overgrowth is often found on plagioclase grains. barite and siderite are important authigenic phases in several intervals, but the presence of barite may be due to the use of drilling mud, potentially contaminating the sand samples with both barium and strontium. most of the sand is fairly loose, but parts of the tyr member are cemented by quartz and calcite as it was located below the oil-water contact, whereas cementation was largely inhibited by oil in most of the kolga member. methods geochemical analyses were performed using a number of methods including inductively coupled plasma mass spectrometry (icp-ms). the advantages of this method compared with x-ray fluorescence (xrf) are that the former measures a wider range of trace elements including rare-earth elements (ree), and that the detection limits are lower than those of xrf, which allows more accurate interpretation of elements found in low concentrations. core material has previously been analysed by friis et al. (2007) using xrf. the modal composition of the sand as well as the chemical composition of the individual mineral grains have been analysed using computer-controlled scanning electron microscopy (ccsem), where each grain is classified as a specific mineral on the basis of its chemical composition (keulen et al. 2008). this method was applied in order to discern whether the samples could be differentiated based on their glauconite composition and to test if injected sand could be identified by its glauconite composition. the >45 μm fraction of the sand was used for the analyses. oil was extracted by toluene, and detergent applied to remove the oily drilling mud and disintegrate the slightly lithified sand. differentiation of palaeogene sand by glauconitic and geochemical fingerprinting, siri canyon, danish north sea mette olivarius, christian knudsen and johan b. svendsen 0° 0° 8°e 58°n 58° germany uk north denmark 250 km norway siri canyon sea 8° fig. 1. map of the north sea region showing the location of the ne–sworiented siri canyon with the tyr and kolga members under investigation in the nini oil field. dashed lines: national borders. © geus, 2011. geological survey of denmark and greenland bulletin 23, 17–20. open access: www.geus.dk/publications/bull 1818 bulk geochemical analyses were carried out at acmelabs, vancouver, on 14 core samples and 73 cuttings samples. major and several minor elements were determined by inductively coupled plasma emission spectrometry (icp-es) on fused glass discs, whereas trace elements, including ree, were identified by icp-ms also on fused glass discs. the modal content of minerals in 10 core samples and 16 samples of the cuttings mounted in epoxy were determined at geus by ccsem on a philips xl40sem (keulen et al. 2008). approximately 1200 grains were analysed per sample. the method integrates backscattered electron micrographs with energy dispersive x-ray spectrometry (edx) to measure the element composition of each grain. the major element weight percentages (wt%) were measured as oxides. the analysis is performed by sweeping over the entire grain, and hence the chemical analysis represents an average of the whole grain and not a point. this is important because the glauconite grains are inhomogeneous. grain size and shape parameters were also measured at the cut surface in the polished section. the fragile nature of the glauconite grains made crushing of the more consolidated parts of the sand inexpedient, so a new application of the ccsem method was developed with measurement of chemical composition in points defined by a grid. this was done in five additional core samples. geochemistry chondrite-normalised ree spectra of the tyr and kolga members are quite similar (fig. 2), except for a positive cerium anomaly in the kolga member. the ree concentrations are moreover higher in the kolga member. the ree spectra of the cuttings samples fit very well with the kolga member (fig. 2d). however, the wells from which the cuttings were sampled have some intervals with high resistivity, and these are generally characterised by a lower content of cerium and an enlarged negative europium anomaly. the lower content of trace elements in the tyr member than in the kolga member makes the sand distinguishable by a number of factors. for example, th and ce in the tyr member are below 6 ppm and 70 ppm, respectively, whereas the concentration is higher in the kolga member. all samples of the cuttings except four are, on this basis, interpreted as kolga member. the four outliers are diluted by either calcite cementation or organic matter, which is seen as high values of calcium and loss on ignition (loi), respectively. glauconite composition the glauconite grains show a wide range in chemical composition, which is reflected in green to brown colours. green grains are usually rounded and well preserved, whereas brown grains show some structural and chemical resemblance to clay minerals. the roundness of the grains could either be caused by their formation process or by subsequent physical abrasion (odin & matter 1981). the best preserved grains are usually those with the highest iron content. zonation seen in many glauconite grains with light centres and dark rims is apparently related to outward decreasing magnesium content. la ce pr nd pm sm eu gd tb dy ho er tm yb lu 1000 100 10 1 100 10 100 10 100 10 sa m p le / r e e c h o n d ri te core samples from the kolga member core samples from the tyr member a b c d cuttings samples interpreted as kolga member kolga member tyr member kolga member (interpreted) fig. 2. ree spectra measured by icp-ms and normalised to the chondrite composition of boynton (1984). a: the tyr member has lower ree concentrations than the kolga member. b: the kolga member is characterised by a small positive cerium anomaly. c: ree spectra from cuttings samples from intervals without infiltration by drilling mud or clayey deposits and with normal resistivity. d: composite diagram with ree spectra from a, b and c indicating that the known and interpreted intervals of the kolga member are idential and that they are different from those of the tyr member. 19 compositional variation is recorded in the glauconite in every sample. however, this variation range is different in the tyr and kolga members. the glauconite in core samples from the tyr member is characterised by a broad scatter and high iron content (fig. 3a), which is distinctly different from the glauconite of the kolga member. the kolga member shows positive correlation between iron and potassium (fig. 3b), which represents a substitution series with aluminium. the kolga member is moreover distinguishable by a high siderite content compared to the tyr member. all the cuttings samples of unknown stratigraphy are interpreted as kolga memcore samples from the kolga member b fe2o3 (wt%) c cuttings samples interpreted as kolga member 0 2 4 6 8 10 12 14 16 18 0 10 20 30 40 50 60 70 80 90 100 core samples from the tyr member a k 2 o ( w t% ) k 2 o ( w t% ) k 2 o ( w t% ) 0 2 4 6 8 10 12 14 16 18 0 10 20 30 40 50 60 70 80 90 100 0 2 4 6 8 10 12 14 16 18 0 10 20 30 40 50 60 70 80 90 100 quartz feldspar barite glauconite carbonate titanite rutile leucoxene ilmenite ti-magnetite chlorite white mica dark mica clinoamphibole/-pyroxene orthoamphibole/-pyroxene siderite pyrite garnet tourmaline zircon fig. 3. fe-k composition of the minerals expressed as fe 2 o 3 versus k 2 o measured by ccsem. a: the tyr member shows a broad glauconite composition without a linear trend. b: the kolga member is characterised by a narrow glauconite composition. c: the sand from the cuttings samples has a glauconite composition that closely resembles that of the kolga member. c u m u la ti ve w t% b c 0 50 100 1 10 100 1000 core samples from the tyr member grain diameter (μm) a 1 10 100 1000 1 10 100 1000 0 50 100 0 50 100 c u m u la ti ve w t% c u m u la ti ve w t% core samples from the kolga member cuttings samples interpreted as kolga member fig. 4. grain-size distribution curves for the minerals measured by ccsem. a: the grain-size distribution of the tyr member has only been measured in one sample, where the glauconite shows a smaller grain size than quartz. b: quartz and glauconite grains in the kolga member are of medium size. c: the cuttings samples have undergone severe crushing and hence the origin of the sand is difficult to determine from the grain-size distribution alone. for legend see fig. 3. 2020 ber on the basis of mineralogy, as their glauconite compositions and siderite contents fit well with this sand (fig. 3c). a large amount of barite is found in many of the cuttings samples, but at least some of it comes from the drilling mud. five of the six samples from the cored tyr member have been measured in single points instead of whole grains because of the extensive cementation, so the results are not entirely reliable. however, the measured glauconite compositions fit well with the broad scatter measured in the un-cemented sample. grain curves the quartz of both the tyr and kolga members is well-sorted, and the variation in grain-size distributions is small (fig. 4). the heavy minerals are finer grained than the light minerals, showing that hydraulic sorting has occurred. the average grain size of the glauconite and quartz grains is almost equal in the kolga member, but the sorting of glauconite is poorer than quartz due to a broad, fine-grained tail, which may be caused by crushing of the fragile glauconite grains. the glauconite in the kolga member is coarser grained than in the tyr member. however, the grain size of the tyr member has only been measured in one sample because of the cementation in the other samples. siderite is silt-sized, and the almost straight grain curves in most samples show that the siderite is authigenic (weibel et al. 2010). cores and cuttings are dominated by quartz grains of about the same size (fig. 4), but the cuttings also contain a fine-grained tail (fig. 4c) which may have been generated by crushing during the drilling process. the glauconite grains are especially susceptible to crushing because of their fragile nature, and this explains why glauconite from cuttings samples is more fine grained than from core samples. concluding remarks the samples of the cuttings collected from the horizontal wells are interpreted as kolga member on the basis of trace element concentrations, ree spectra, glauconite compositions and siderite contents. this implies that remobilisation is restricted to intra-strata processes, rather than between strata. modelling of the injected part of the field is therefore likely to be comparable to that of the in situ parts, as the original composition of the sand is the same. icp-ms and ccsem have proved useful in characterising sand types, and from these observations it was possible to identify the origin of the intrusive sand bodies. especially the ree spectra measured by icp-ms and the glauconite compositions measured by ccsem have enhanced the understanding of the sediments. acknowledgements this study was conducted in cooperation with the partnership of licence 4/95 in the danish north sea, operated by dong energy. the partnership is thanked for permission to publish the results. references boynton, w.v. 1984: geochemistry of the rare earth elements: meteorite studies. in: henderson, p. (ed.): rare earth element geochemistry, 63–114. amsterdam: elsevier. friis, h., poulsen, m.l.k., svendsen, j.b. & hamberg, l. 2007: discrimination of density flow deposits using elemental geochemistry – implications for subtle provenance differentiation in a narrow submarine canyon, palaeogene, danish north sea. marine and petroleum geology 24, 221–235. hamberg, l., dam, g., wilhelmson, c. & ottesen, t.g. 2005: paleocene deep-marine sandstone plays in the siri canyon offshore denmark, southern norway. in: doré, a.g. & vining, b.a. (eds): petroleum geology: north-west europe and global perspectives, 1185–1198. proceedings of the 6th petroleum geology conference. london: geological society. keulen, n., frei, d., bernstein, s., hutchison, m.t., knudsen, c. & jensen, l. 2008: fully automated analysis of grain chemistry, size and morphology by ccsem: examples from cement production and diamond exploration. geological survey of denmark and greenland bulletin 15, 93–96. odin, g.s. & matter, a. 1981: de glauconiarum origine. sedimentology 28, 611–641. poulsen, m.l.k., friis, h., svendsen, j.b., jensen, c.b. & bruhn, r. 2007: the application of bulk rock geochemistry to reveal heavy mineral sorting and flow units in thick, massive gravity flow deposits, siri canyon palaeocene sandstones, danish north sea. developments in sedimentology 58, 1099–1121. schiøler, p. et al. 2007: lithostratigraphy of the palaeogene – lower neogene succession of the danish north sea.  geological survey of denmark and greenland bulletin 12, 77 pp. weibel, r., friis, h., kazerouni, a.m., svendsen, j.b., stokkendal, j. & poulsen, m.l.k. 2010: development of early diagenetic silica and quartz morphologies – examples from the siri canyon, danish north sea. sedimentary geology 228, 151–170. authors’ addresses m.o. & c.k., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: mol@geus.dk j.b.s., dong energy, exploration and production, agern allé 24 –26, dk-2970 hørsholm, denmark. geological survey of denmark and greenland bulletin 41, 2018, 75-78 75 initial field activities of the camp century climate monitoring programme in greenland william colgan, allan pedersen, daniel binder, horst machguth, jakob abermann and mike jayred fig. 1. location of camp century that was constructed by the us army corps of engineers in 1959 in the greenland ice sheet. camp century was abandonned in 1967. 500 km camp century thule air base camp century was a military base constructed by the us army corps of engineers in 1959 in the near-surface layers of the greenland ice sheet at 77.13°n, 61.03°w and 1886 m above sea level (clark 1965; fig. 1). the base housed up to 200 military personel and was continuously occupied until 1964. after three years of additional seasonal operation, the base was abandoned with minimal decommissioning in 1967. recent danish scholarship has documented the political and military history of camp century in detail (nielsen & nielsen 2016). in 2016, the geological survey of denmark and greenland (geus) participated in a multi-nation study that presented regional climate simulations that suggested the icesheet surface mass balance at camp century may change from net accumulation to net ablation by 2100 under the un intergovernmental panel on climate change rcp8.5 ‘business-as-usual’ climate scenario. however, according to colgan et al. (2016), net accumulation would persist beyond 2100 at camp century under the climate-change mitigation characterised by rcp4.5, an approximately ‘paris agreement’ climate scenario. in 2017, in response to concerns from the government of greenland over the potential to remobilisation of contaminants from camp century within the next century, the government of denmark established a programme for long-term climate monitoring and detailed one-time surveying of the debris field at camp century (colgan et al. 2017). this report describes the initial field activities of the camp century climate monitoring programme in the context of the four programme goals: 1. to continuously monitor relevant climate variables, including the depth to which meltwater percolates, at the camp century site. 2. to regularly update annual likelihoods of meltwater interacting with abandoned materials at the camp century site over the next century. 3. to map the estimated spatial extent and vertical depth of abandoned wastes across the camp century site. 4. to publicly report all findings from the camp century climate monitoring programme in a timely manner. field logistics field activities of the camp century climate monitoring programme were initiated in summer 2017, when a six-person team spent two weeks at the camp century site (19 july to 3 august). there is no abandoned infrastructure visible at the ice-sheet surface at camp century (fig. 2). debris field location, as well as zones of restricted drilling depth, were estimated prior to field work (fig. 3). this was done by georeferencing a historical site map using a single tie-point, the location of the original drill tower, corrected for motion since its last precise survey in 1986 (gundestrup et al. 1987). the 2017 summer camp, which consisted of three common © 2018 geus. geological survey of denmark and greenland bulletin 41, 75–78. open access: www.geus.dk/bulletin http://www.geus.dk/bulletin 7676 tents and six personal tents, was established within the debris field, approximately aligned with the location of post-closure summer camps (kovacs 1970). although decamped entirely, the footprint of the 2017 summer camp will likely be visible in subsequent ice-penetrating radar surveys due to the formation of massive wind-sculpted snowdrifts around it. during field work, there was persistent cloud cover with frequent storm conditions. the mean wind speed was 8.8 m/s, and the maximum 1 hour mean wind speed was 18.3 m/s (beaufort 8). thule air base, located approximately 200 km west, served as logistical base for the field work. a ski-equipped twin otter aircraft was used to transport 3200 kg of equipment and supplies to and from the ice sheet. field work consisted of installing three automated instrument stations, drilling boreholes for instrument installation and firn sampling, surveying velocity stakes, and collecting ice-penetrating radar profiles. of the 175 m firn core drilled, 135 m were analysed in the field and 40 m were transported to copenhagen for more detailed radionuclide analysis by the center for nuclear technologies at the technical university of denmark. instruments and data climate measurements were initiated using automated weather station technology previously developed by geus. the automatic weather station design has a proven record of more than 175 station-years of deployment in greenland since its introduction in 2007 (citterio et al. 2015). the primary weather station at camp century (cen) measures air temperature and humidity, wind speed and direction, atmospheric pressure, upward and downward shortwave and longwave radiation, subsurface (snow/ice) temperatures to 10 m s6 s5 cen-com b62 500 m b73 cen cen-gps cen-thm s7 instrument tower borehole ice velocity stake 2017 camp ice-penetrating radar 25 m drill zone 10 m drill zone 545250 545750 546250 546750 547250 547750 548250 85 66 50 0 85 67 00 0 85 67 50 0 85 68 00 0 85 68 50 0 85 69 00 0 projection: wgs84, utm 20n s2 s3 s4 fig. 3. overview of initial field activities at camp century. recently constructed instrument towers refer to the primary weather station (‘cen’), the supplementary thermistor station (cen-thm), the supplementary compaction station (cen-com) and the supplementary global positioning system station (cen-gps; to be installed). boreholes refer to the 73 (b73) and 62 (b62) m firn cores. restricted drill zone depths were assessed based on georeferencing of a historical as-built site map (kovacs 1970). fig. 2. the temporary ice-sheet camp at 1600 utc on 20 july 2017. the camp consisted of three common tents and six personal tents. there was persistent cloud cover with frequent storm conditions and a mean wind speed of 8.8 m/s. no abandoned infrastructure is visible at the ice-sheet surface; the entire debris field is now subsurface as a result of net snow accumulation since closure. 77 fig. 4. left: near-surface firn density profiles measured both inside (b73) and outside (b62) the debris field. the high-density layer between 32 and 35 m depth within the debris field likely reflects enhanced compaction during the c. 1960–1964 active period. right: near-surface snow/ice temperatures at b73 inside the debris field. the annual temperature cycle penetrates to c. 12 m depth, with year-round firn temperatures remaining c. –24°c below this depth. density (kg/m )3 400 600 800 d ep th (m ) 0 10 20 30 40 50 60 70 b62 b73 temperature (°c) –30 –20 –10 0 0 10 20 30 40 50 60 70 aug sep oct nov dec jan feb mar apr depth, and snow depth, as well as diagnostic parameters such as battery voltage. the temperature of the relatively porous near-surface ice-sheet layer known as firn is also measured by a supplementary thermistor station (cen-thm) to a depth of 73 m, which approximates local pore close-off depth (fig. 3). ice-sheet structure was analysed with 73 and 62 m deep firn cores, hereafter referred to as b73 and b62. firn density and melt percentage were measured to a minimum depth of 62 m at locations inside and outside the debris field (fig. 4). a third automatic station was deployed to measure the compaction rate of snow into ice, or vertical strain, over the 0 to 5 m, 0 to 20 m and 1 to 62 m depth ranges. this third automatic station, the supplementary compaction station (cen-com), is located outside the debris field. all three automatic stations satellite-transmit their measurements in near-real-time to www.campcenturyclimate.dk. a fourth station that records observations from the global positioning system (cen-gps) will be installed to continuously monitor ice flow. these climate and ice data will be used to calibrate and validate future simulations of firn evolution. preliminary interpretation the measurements from the automated weather station (cen) record that midday air temperature exceeded 0°c for three days during the operational period of the 2017 summer camp. the maximum one-hour mean air temperature was 1.8°c. under these warm conditions, surface melt quickly froze to moving drill parts when the drill was lowered into cold winter firn. this necessitated a two-day suspension of drilling. preliminary analysis of the b73 deep thermistor measurements, located within the debris field, indicates that the annual temperature cycle in near-surface snow and ice temperatures penetrates to c. 12 m depth (fig. 4). beyond this depth, year-round firn temperatures appear to remain c. –24°c. in summer 2017, there was limited meltwater production and refreezing, with no apparent change in firn temperature beyond this annual diffusion cycle. preliminary analysis of near-surface firn structure indicates that refrozen meltwater layers are readily identifiable in the uppermost 15 m of the firn. the largest of these layers is c. 8 cm thick, which represents the melt-and-refreeze of c. 25% of annual snowfall (buchardt et al. 2012). the firn cores suggest that meltwater movement beyond the annual layer is unlikely. near-surface firn densities are similar both inside and outside the debris field to 32 m depth (fig. 4). between 32 and 35 m depth, firn density is significantly greater within the debris field. this high-density layer is slightly discoloured in appearance and likely reflects enhanced compaction and pollution during the c. 1960–1964 active period. below this active layer, firn densities are similar inside and outside the debris field. after drilling b73 through the highdensity active layer, pressurised hydrocarbon vapours vented from the borehole until it was backfilled. mobile hydrocarhttp://www.campcenturyclimate.dk 7878 bon vapours were not anticipated, and vapour-tight equipment was not available on-site to opportunistically sample these vapours. radar survey data from the one-time summer 2017 radar survey are being analysed to perform a detailed assessment of the horizontal extent and vertical range of the debris field. 100 and 250 mhz ice-penetrating radar data were collected by crosscountry skiing in a dense grid pattern over the camp century debris field (fig. 3). the radar profiles, each tagged with global positioning system coordinates, will be available on the programme website. a preliminary field analysis of the ice-penetrating radar data shows that the sub-surface debris field is c. 2 km in diameter, with debris ranging between c. 20 and 100 m depth. this ice-penetrating radar data will permit improved geo-referencing of historical as-built site maps, via precisely positioning key subsurface infrastructure features, which will facilitate delineating the debris field beyond the extent recorded by as-built site maps. programme outlook this report describes the initial field activities of the camp century climate monitoring programme in the context of programme goals. near-real-time climate and ice measurements from automated stations, ice-penetrating radar profiles, as well as programme outreach materials and publications, can be accessed at www.campcenturyclimate.dk. subsequent field work at camp century will be undertaken, as needed, to service deployed instrumentation. during these subsequent site revisits, ice-velocity stakes will be resurveyed to precisely measure the relatively slow (<5 m/yr) ice velocity over several years. data analysis, in support of observationally-constrained numerical simulations of the evolution of meltwater and firn, is the major programme focus. while climate change now gives camp century previously unanticipated social significance, the sustained effort of the camp century climate monitoring programme will continue to provide danish and greenlandic stakeholders open access to relevant in situ measurements and model projections. refined knowledge of the spatial extent and vertical range of the debris field, as well as the changes in firn structure and meltwater production anticipated under climate change, will inform science-based discussions of the shifting fate of camp century. acknowledgements the camp century climate monitoring programme is jointly funded by geus and the danish cooperation for environment in the arctic (dancea) within the danish ministry for energy, utilities and climate. j.a. was supported by the greenlandic ministry of independence, foreign affairs and agriculture. the 2017 field team (w.c., a.p., d.b., h.m., j.a. and m.j.) warmly thank danish liaison officers kim marchuard mikkelsen and jens alsing for logistical assistance at thule air base. references buchardt, s., clausen, h., vinther, b. & dahl-jensen, d. 2012: investigating the past and recent δ18o-accumulation relationship seen in greenland ice cores. climate of the past 8, 2053–2059. clark, e.f. 1965: camp century evolution of concept and history of design, construction and performance. cold regions research and engineering laboratory. technical report 174, 69 pp. citterio, m. et al. 2015: automatic weather stations for basic and applied glaciological research. geological survey of denmark and greenland bulletin 33, 69–72. colgan, w., machguth, h., macferrin, m., colgan, j.d., van as, d. & macgregor, j.a. 2016: the abandoned ice sheet base at camp century, greenland, in a warming climate. geophysical research letters 43, 8091–8096. colgan, w., andersen, s.b., van as, d., box, j.e. & gregersen, s. 2017: new programme for climate monitoring at camp century, greenland. geological survey of denmark and greenland bulletin 38, 57–60. gundestrup, n.s., clausen, h.b., hansen, b.l. & rand, j. 1987: camp century survey 1986. cold regions science and technology 14, 281– 288. kovacs, a. 1970: camp century a pictorial overview june 1969. cold regions research and engineering laboratory, special report 150, 59 pp. nielsen, h. & nielsen, k. 2016: camp century – cold war city under the ice. in: doel, r., harper, k. & heymann, m. (eds): exploring greenland: cold war science and technology on ice, 195–216. new york: palgrave macmillan us. authors’ addresses w.c., a.p., d.b., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: wic@geus.dk. h.m., university of fribourg, avenue de l’europe 20, sx 1700 fribourg, switzerland. j.a., asiaq greenland survey, qatserisut 8, dk 3900 nuuk, greenland. m.j., university of wisconsin, madison, wisconsin 53706, usa. http://www.campcenturyclimate.dk mailto:wic@geus.dk geological survey of denmark and greenland bulletin 41, 2018, 57-62 57 the west greenland continental margin has been the subject of petroleum exploration by companies and research projects since the 1970s and many data have been acquired since. licensing rounds issued by the greenland authorities in 2002 and 2004 offshore southern west greenland resulted in company licenses which led to data acquisition and three exploration wells. the extensive new data form a basis for updated mapping by means of data, new analyses of the subsurface geology and improved understanding of the stratigraphy and the geological development. the geological survey of denmark and greenland (geus) has recently completed a comprehensive mapping project of the subsurface in an area covering 116 000 km2 offshore southern west greenland (fig. 1). the results include maps displaying large structural highs and faults, cretaceous sedimentary basins and volcanic areas, illustrated by cross-sections through the area. a new seismic stratigraphy with eight mega-units from the seabed to the basement was also defined. in addition, studies from wells of biostratigraphy and petrology were carried out that provide important new information. the new data include extensive 2d seismic data and eight wells including the three exploration wells at2-1, at7-1 and lf7-1 drilled in 2011 by cairn energy (fig. 1). key results of the work are summarised below. geological setting and scope of project the southern west greenland continental margin is located between west greenland to the east and the oceanic crust and parts of the davis strait high to the west (fig. 2). a number of rifted basins with large structural highs are interpreted to have developed offshore southern west greenland during the cretaceous (chalmers et al. 1993; sørensen 2006; gregersen 2014). cretaceous and cenozoic sedimentary successions were previously defined from wells and outcrops on the central and southern west greenland continental margin (rolle 1985; dam et al. 2009; nøhr-hansen et al. 2016). during the paleocene and eocene, oceanic crust developed between canada and greenland, and the cretaceous rifted continental margin of west greenland was separated from eastern canada (oakey & chalmers 2012). the large-scale movements between greenland and canada generated new structures during the palaeogene and reactivated faults within cretaceous basins. the purpose of the research-based project described here was to update previous subsurface mapping with the most recent data to provide an improved knowledge of the structures and the geological development, which can lead to an evaluation of the resource potential. the project was carried out in 2015–2017 for the ministry of mineral resources in nuuk and incorporated a number of sub-tasks including: (1) seismic interpretation and mapping; (2) well correlation; (3) biostratigraphy; (4) analyses of potential reservoir rocks in wells from 2011 and (5) analyses of igneous rocks, basement and provenance. this paper describes some of the key results from the seismic interpretation, mapping and well data-renew subsurface mapping offshore southern west greenland using geophysical and geological data ulrik gregersen, morten s. andersen, henrik nøhr-hansen, emma sheldon, thomas f. kokfelt, mette olivarius, christian knudsen, kristian g. jakobsen and jan s. adolfssen fig. 1. map of the study area offshore southern west greenland with positions of 2d seismic lines and exploration wells. the positions of figs 3 and 4 are also shown. seismic line well 100 km 58°w 56°w 54°w 52°w 65°n 58°w 52°w 66°n 67°n 64°n 63 °n 65°n 58°w 56°w 54°w 52°w 58°w 52°w 66°n 67°n 64°n 63 °n greenland ikermiut-1 kangâmiut-1 nukik-2 nukik-1 at7-1 at2-1 nuuk fig. 3 fig. 4 lf7-1 qulleq-1 c an ad a greenland ikermiut-1 kangâmiut-1 nukik-2 nukik-1 at7-1 at2-1 nuuk fig. 3 fig. 4 lf7-1 qulleq-1 c an ad a © 2018 geus. geological survey of denmark and greenland bulletin 41, 57–62. open access: www.geus.dk/bulletin http://www.geus.dk/bulletin 5858 lated tasks. the lateral extent of the main tectonic elements, including structural highs, faults and basins, is outlined in a structural elements map (fig. 2). methods interpretation of seismic stratigraphic horizons and units was carried out using schlumberger petrel © software and included data from wells and seismic surveys (fig. 1). in addition, other data including gravity and magnetic surveys and seabed sampling were used for the interpretation and mapping. most structural highs and basins are defined and outlined by seismic interpretation in combination with gravityand magnetic anomaly maps, and are supported by other data such as published refraction models. a number of methods were used at geus for the other studies, including scanning electron microscopy, porosity & permeability measurements, and rock/mineral analyses and radiometric dating using a laser ablation inductively coupled plasma mass spectrometer. biostratigraphic dating based on palynology and micropalaeontology was carried out on a total of more than 200 samples from the at2-1, at-7-1 and lf7-1 wells. seismic stratigraphy and structures a robust stratigraphic framework was established with eight seismic stratigraphic mega-units (a–h) divided by horizons from the seabed (a1) to the top of pre-cretaceous sedimentary rocks or the acoustic basement (h1) shown in fig. 3. the mega-units include internal tops of sub-units (e2, ev, f2, fv and hx) described below, and new biostratigraphy was used to constrain ages of the units. the units and most horizons are shown in seismic cross-sections across deep rifted basins separated by large structural highs (figs 3, 4). most of the structural highs, faults and basins trend se–nw but a few strike in a more northerly direction or towards ne (fig. 2). the total succession between the seabed and the basement has been mapped and shows thick basins between the main structures (fig. 5), and also more local structures where the wells were drilled. some basins and parts of basins are untested by wells. mega-unit h; pre-cretaceous basement mega-unit h includes the lower parts of basins and the basement of pre-cretaceous ages (fig. 3). it mostly includes the acoustic basement in large structures below the cretaceous and cenozoic basins, and was drilled in the at7-1 and lf7-1 wells (figs 3, 4). geus’ analyses of igneous rock samples from lower parts of the at7-1 well mainly reveal granites, granodiorites and tonalities. u-pb dating of zircons from some of the samples gives ages of c. 2730–3190 ma. the samples are from the lower part of mega-unit g and the uppermost part of mega-unit h in the drilled structure (fig. 3) in a succession fig. 2. structural elements map of the southern west greenland continental margin. the map shows areas with structural highs, faults, cretaceous basins and palaeogene volcanic cover. the majority of the large structures and faults trends towards the nw with a few towards the ne. in addition, the positions of the wells and the boundary to canada are shown. fsc: fylla structural complex with its composite structures, incl. fcb: fylla central basin. lfh: lady franklin high. mb: maniitsoq basin. kr: kangâmiut ridge. ifz: ikermiut fault zone. structural element names are mainly from chalmers et al. (1993), sørensen (2006) and døssing (2011). the geological map of west greenland onshore is from henriksen et al. (2009). the position of the oceanic crust and the davis strait high in the canadian sector are from oakey & chalmers (2012). the names of the wells are shown in fig. 1. n ukik platformkr ifz kang âm iut ba sin sis im iut ba sin maniitsoq high hecla high mb fylla east basin fsc south hecla highlfh la dy f ra nk lin b as in paamiut basin fy lla w es t h igh fcb nuuk basin greenland d av is s tr ait h igh c an ad a g re en lan d l a b r a d o r s e a nuuk (d sh ) 66°n 67°n 65°n 64°n 63°n 60°w 54°w 52°w 50°w56°w58°w extensional fault compressional/ transtensional fault well proterozoic basement oceanic crust structural highs cretaceous basins palaeogene volcanics, partly covering cretaceous basins archean basement 100 km 59 with interbedded sedimentary layers. similar igneous rocks, some of which have been weathered, were also reported from the succession by cairn energy (2011). the rocks are mostly similar in age and composition to parts of the basement terrane in southern west greenland. metamorphic rocks reported by cairn energy (2011) from the lowermost part of the lf7-1 well occur just below the h1 horizon. however, pre-cretaceous sedimentary rocks are also expected to be locally preserved in basins such as in the sisimiut basin (fig. 2), where a sub-unit is present in the upper part of mega-unit h. ordovician marine carbonates were sampled from the seabed over the davis strait high, where some possibly are in situ (stouge et al. 2007). such rocks may also occur in the adjacent basins. in addition, an organic-rich ordovician sample from the davis strait was considered to have source rock potential (bojesen-koefoed 2011). cretaceous and paleocene oil seeps have been described from the nuussuaq basin farther north with a possible wider occurrence in west greenland (bojesen-koefoed 2011). mega-unit g; early to mid-cretaceous large rift structures, local wedge-shaped units within megaunit g and extensional faults developed during the early to mid-cretaceous (figs 3, 4). bowl or v-shaped, strong reflections near the g1 horizon (fig. 3) are interpreted as sills of a larger sill complex in the nuuk basin and lady franklin basin. mega-unit g is most likely late albian to cenomanian in age. a seismic correlation (fig. 3) with the at7-1 well shows that the g1 surface divides a cenomanian–turonian sedimentary succession. the overlying lower part of mega-unit f is dominated by sandstones with conglomerates and thin claystones. the underlying mega-unit g includes a thick succession dominated by conglomerates and a lowermost ?late albian – early cenomanian succession including rocks from an igneous basement, interbedded thin sandstones, conglomerates and claystones. lithologies in the at2-1, at71 and lf7-1 wells were determined by cairn energy (2011) and in this study. mega-unit f; late cretaceous – paleocene the present study shows that the lower succession of megaunit f is mainly of cenomanian–turonian age. variable reservoir properties were found from the study of well samples. in the lower part of mega-unit f of the at7-1 well, the analysed conglomerate samples show intermediate to good reservoir quality, e.g. a sidewall core sample with 24.77% porosity and 78.73 md permeability. the sandstones and overlying diatomite have high porosities, and medium to low and low permeability, respectively. the other cretaceous rocks, analysed in the other two wells (at2-1 and lf7-1), mostly show poorer reservoir quality. however, sandstones with reservoir potential are expected to be present in some of the structures. a 323 m thick succession dominated by sandstones of late santonian age with reservoir potential is described from the lower part of the qulleq-1 well in the fylla structural complex (christiansen et al. 2001; fig. 2) and its top is also mapped in this study. during parts of the late cretaceous, relative tectonic quiescence prevailed in most areas and thick claystone-dominated units were deposited in the basins. this study shows that parts of mega-unit f (figs 3, 4) are dominated by thick upper cretaceous to paleocene successions, which consist mostly of claystones. such successions are drilled in the lf7-1, at2-1, qulleq-1 and ikermiut-1 wells, which include the ikermiut fm (rolle 1985). however, tectonic movements and erosion probably related to rifting and/or uplift may have occurred in the study area during late cretaceous to paleocene, in some places indicated by truncation and faulting of structures (see also chalmers et al. 1993). erosion or non-deposition may also be indicated by missing sections in wells. hiati of different timespans seem to occur during the late cretaceous to early paleocene (mostly from within the late campanian to danian) in the at2-1 and at7-1 wells (in at7-1 even to the coniacian), and were also previously noted in the qulleq-1 and ikermiut-1 wells (nøhr-hansen et al. 2016). volcanism the lower parts of the at2-1 well succession contain volcanic rocks and thin claystones, which are biostratigraphically dated in the present study as late cenomanian to early turonian. the horizon fv correlates with the top of the volcanic succession which may partly cross or overlie mega-unit g (fig. 3). during the paleocene–eocene, flood basalts and other volcanic rocks were deposited in large parts of the west greenland continental margin (skaarup 2001; sørensen 2006; larsen et al. 2016). the tops of extensive volcanic areas and highs are mapped offshore southern west greenland (fig. 2) primarily from seismic reflection data at horizon ev (fig. 4) and from magnetic anomaly data. these volcanic areas include the hecla high, the maniitsoq high, the davis strait high and parts of the nukik high (fig. 2). the nukik-2 well on the nukik high includes a lower succession with hyaloclastite beds and thicker intrusives (dolerites), where the upper part has been biostratigraphically dated as late paleocene (hald & larsen 1987). an almost 700 m thick succession with paleocene subaerial basaltic lava flows was drilled in the lower 6060 part of the hellefisk-1 well (hald & larsen 1987) north of the study area. paleocene basalts were sampled from the seabed over the davis strait high and south of the hecla high, where additional early eocene basalts were recovered (larsen & dalhoff 2006). a refraction seismic study by funck et al. (2007) also shows basalts in the western parts of the study area and farther west with a tie to the canadian gjoa g-37 well. this well includes basalts with thinner mudstones in a >1 km thick danian to thanetian interval (nøhr-hansen et al. 2016). the palaeogene volcanic rocks mostly cover parts of structures and cretaceous sedimentary basins in the study area (figs 2, 4), in parts of the davis strait – southern baffin bay areas (gregersen & bidstrup 2008) and in the nuussuaq basin (dam et al. 2009; larsen et al. 2016). the top of the volcanic successions (ev) is overlain by the seismic horizon e2 of late paleocene age (late thanetian from nøhr-hansen et al. 2016) towards the south (fig. 4). geological development of younger units the west greenland continental margin moved towards ne and n in connection with the palaeogene sea-floor spreading between canada and greenland (oakey & chalmers 2012). these movements caused compression–transtension tectonism with thrust faults and associated basins along the davis strait high and the ikermiut fault zone during the late paleocene to eocene (fig. 2; gregersen & bidstrup 2008). seismic geometries in mega-units e and d with irregular sub-units near faults and basin mounds suggest mass-flows including slumps and basin fans (fig. 3). they were formed during mainly the eocene and miocene and may include potential leads for hydrocarbon. parts of the mass-flows may be related to movements during the formation of palaeogene oceanic crust. parts of upper eocene, oligocene and lower miocene successions are mostly absent in wells from the area (nøhr-hansen et al. 2016). the upper parts of mega-unit d and mega-unit c are miocene in age, mega-unit b is possibly late miocene to pliocene and mega-unit a is probably late pliocene to pleistocene in age (fig. 3). the ages of the mega-units are indicated from biostratigraphy (piasecki 2003; nøhr-hansen et al. 2016). in miocene to pleistocene successions, fewer large faults and other indications of tectonism are found. large contourite drifts, slides and other forms of mass-mobilisation developed during the miocene to pleistocene (nielsen et al. 2011) within mega-units a–c (fig. 4). in addition, sand sw-directed shelf progradation occurred, fig. 3. composite nw–se seismic section across the nuuk basin with large structures and the at2-1 and at7-1 wells. the succession from the seabed to the basement is divided by seismic stratigraphic horizons a1–h1 into seismic mega-units a–h and tentative ages (maximum timespans) of major successions are shown in colour. the deeper parts of the section are dominated by a cretaceous sedimentary succession with rifted basins, and with a local volcanic succession and sills. the seismic lines shown (tgs-green2003-29, tgs-gr2000-215 and enc2005-1) are courtesy of tgs-nopec geophysical company asa. the location of the section is shown in fig. 1. nw seat2-1at7-1 0 tw ow ay tr av el ti m e (s ec .) 1 2 3 4 5 nuuk basin a1 b1 c1 d1 e1 e2 f1 g1 sills h1 a b c d e f fv v v v g h basement fy lla b ou nd ar y fau lt acoustic basement cretaceous to paleocene well fault v: near top of volcanic layers late cenozoic early to mid-cenozoic 10 km 61 probably as a result of glaciations, and a near-horizontal succession also formed in mega-unit a. conclusions a study was carried out with subsurface mapping using geophysical and well data offshore southern west greenland. a seismic stratigraphy with eight mega-units (a–h) from the seabed to the pre-cretaceous basement has been defined in the area and shown on cross-sections. a new structural elements map displaying the main structures, basins and faults, as well as a sedimentary thickness map are presented. fig. 4. the lf7-1 well and a composite nw–se seismic section across the lady franklin basin. the succession from the seabed to the basement is divided by seismic stratigraphic horizons a1–h1 into seismic mega-units a–h and tentative ages of major successions are shown in colour (fig. 3). it is dominated by cretaceous successions of mega-units f and g and includes rifted structures. mega-unit f is overlain by a probably thin palaeogene volcanic succession. a miocene–pliocene contourite succession of mega-units b and c occurs south-east of the lf7-1 well. the seismic line shown (tgs-blf2005-43) is courtesy of tgs-nopec geophysical company asa. the location of the section is shown in fig. 1. fig. 5. thickness isochore map in two-way travel time (ms) between the seabed (a1 horizon) and the pre-cretaceous basement (h1 horizon). the map includes mostly sedimentary successions but locally also includes volcanic successions as illustrated in figs 2–4. thick dominantly sedimentary successions occur in mainly the lady franklin basin, the fylla east basin, the nuuk basin, the kangâmiut basin and the sisimiut basin, whereas successions thin over structural highs (figs 2–4). the names of the wells are shown in fig. 1. 0 1 2 3 4 5 nw lf7-1 se lady franklin basin lady franklin basinhecla high lfb hecla high a1 b1 d1 e2 ev a b c d e f g h e1 f1 c1 g1 f1x h1 10 km basement tw ow ay tr av el ti m e (s ec .) 600 1200 1800 2400 3000 3600 4200 4800 thickness time (ms) 52°w58°w 58°w 66°n 67°n 54°w56°w 63ºn 64°n nuuknuuk 65°n 100 km 6262 biostratigraphic ages, radiometric dating and lithologies from well data are correlated to parts of the mega-units. the studied wells comprise cretaceous to cenozoic sedimentary successions, a volcanic succession and granitic rocks from the basement. the results presented here are based on analyses from this study and recently released data from the wells drilled in 2011 by cairn energy (at2-1, at7-1 and lf7-1). the results have improved the understanding and outline of the large-scale structures, basins and provide input for further work in the region including new resource-potential evaluations, but also point out uncertainties and risks that require clarification. acknowledgements the study was co-financed by the ministry of mineral resources (government of greenland) and geus. the constructive comments from the referees kate dickie and christopher harrison are much appreciated. the displayed seismic lines are courtesy of tgs-nopec geophysical company asa. references bojesen-koefoed, j.a. 2011: west greenland petroleum systems. an overview of source rocks and oil seepages and their implications for offshore petroleum exploration. danmarks og grønlands geologiske undersøgelse rapport 2011/42, 49 pp. cairn energy 2011: licensing reporting of cairn energy including the geological end of well reports and the composite well-logs from the at2-1, at7-1 and lf7-1 exploration wells. chalmers, j.a., pulvertaft, t.c.r., christiansen, f.g., larsen, h.c., laursen, k.h. & ottesen, t.g. 1993: the southern west greenland continental margin: rifting history, basin development, and petroleum potential. in: parker, j.r. (ed.): petroleum geology of nw europe: proceedings of the 4th conference. geological society, london, 915–931. christiansen, f.g., bojesen-koefoed, j.a., chalmers, j.a., dalhoff, f., mathiesen, a., sønderholm, m., dam, g., gregersen, u., marcussen, c., nøhr-hansen, h., piasecki, s., preuss, t., pulvertaft, c.r., rasmussen, j.a. & sheldon, e. 2001: petroleum geological activities in west greenland in 2000. geology of greenland survey bulletin 189, 24–33. dam, g., pedersen, g.k., sønderholm, m., midtgaard, h., larsen, l.m., nøhr-hansen, h. & pedersen, a.k. 2009: lithostratigraphy of the cretaceous-paleocene nuussuaq group, nuussuaq basin, west greenland. geological survey of denmark and greenland bulletin 19, 171 pp. døssing, a. 2011: fylla bank: structure and evolution of a normal-toshear rifted margin in the northern labrador sea. geophysical journal international 187, 655–676. funck, t., jackson, h.r., louden, k.e. & klingelhöfer, f. 2007: seismic study of the transform-rifted margin in davis strait between baffin island (canada) and greenland: what happens when a plume meets a transform. journal of geophysical research 112, 1–22. gregersen, u. 2014: the west greenland continental margin. in: hopper, j.r. et al. (eds): tectonostratigraphic atlas of the north-east atlantic region. copenhagen: geological survey of denmark and greenland, 321–325. gregersen, u. & bidstrup, t. 2008: structures and hydrocarbon prospectivity in the northern davis strait area, offshore west greenland. petroleum geoscience 14, 151–166. hald, n. & larsen, j.g. 1987: early tertiary, low-potassium tholeiites from exploration wells on the west greenland shelf. grønlands geologiske undersøgelse rapport 136, 25 pp. henriksen, n., higgins, a.k., kalsbeek, f. & pulvertaft, t.c. 2009: greenland from archaean to quaternary – descriptive text to the 1995 geological map of greenland 1:2 500 000. 2nd edition. geological survey of denmark and greenland bulletin 18, 126 pp. larsen, l.m. & dalhoff, f. 2006: composition, age, and geological and geotectonic significance of igneous rocks dredged from the northern labrador sea and the davis strait. danmarks og grønlands geologiske undersøgelse rapport 2006/43, 67 pp. larsen, l.m., pedersen, a.k., tegner, c., duncan, r.a., hald, n. & larsen, j.g. 2016: age of tertiary volcanic rocks on the west greenland continental margin: volcanic evolution and event correlation to other parts of the north atlantic igneous province. geological magazine 153, 487–511. nielsen, t., andersen, c., knutz, p.c. & kuijpers, a. 2011: the middle miocene to recent davis strait drift complex: implications for arctic–atlantic water exchange. geo-marine letters 31, 419–426. nøhr-hansen, h., williams, g.l. & fensome, r.a. 2016: biostratigraphic correlation of the western and eastern margins of the labrador-baffin seaway and implications for the regional geology. geological survey of denmark and greenland bulletin 37, 74 pp. oakey, g.n. & chalmers, j.a. 2012: a new model for the palaeogene motion of greenland relative to north america: plate reconstructions of the davis strait and nares strait regions between canada and greenland. journal of geophysical research 117, 1–28. piasecki, s. 2003: neogene dinoflagellate cysts from davis strait, offshore west greenland. marine and petroleum geology 20, 1075–1088. rolle, f. 1985: late cretaceous–tertiary sediments offshore central west greenland: lithostratigraphy, sedimentary evolution, and petroleum potential. canadian journal of earth science 22, 1001–1019. skaarup, n. 2001: offshore volcanic rocks in baffin bay. a seismic interpretation of the structures and development of the palaeogene offshore volcanic rocks in central west greenland and on the baffin island margin eastern canada. phd thesis from univerity of copenhagen. danmarks og grønlands geologiske undersøgelse rapport 2001/117, 154 pp. sørensen, a.b. 2006: stratigraphy, structure and petroleum potential of the lady franklin and maniitsoq basins, offshore southern west greenland. petroleum geoscience 12, 221–234. stouge, s., ineson, j.r., rasmussen, j.a. & dalhoff, f. 2007: sedimentary dredge samples from the davis strait high: stratigraphic and palaeoenvironmental implications. danmarks og grønlands geologiske undersøgelse rapport 2007/19, 49 pp. authors’ addresses u.g., m.s.a., h.n.-h., e.s., t.f.k., m.o. & c.k.,geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ug@geus.dk. k.g.j. & j.s.a., ministry of mineral resources, government of greenland (nuuk), p.o. box 930, dk-3900 nuuk, greenland. mailto:ug@geus.dk geological survey of denmark and greenland bulletin 4, 2003, pp 53-56 53 flood basalt-covered basins exist worldwide along continental margins and are now in focus as targets for future hydrocarbon exploration. it is generally difficult to image through the basalt cover by conventional seismic reflection methods, and this is a major challenge to future petroleum exploration offshore the faroe islands. long-offset profiling has proven very successful (white et al. 2003). surprisingly, however, it is possible to image through kilometre-thick basalt sequences on some conventional profiles. details of basalt stratigraphy are revealed on old, reprocessed seismic profiles as well as on recently acquired profiles, even though the imaging may be unsuccessful on nearby profiles (e.g. boldreel & andersen 1993). this stresses the need for a better understanding of the acoustic and other physical properties of basalt as well as of the degree of three-dimensional heterogeneity. the seifaba project (seismic and petrophysical properties of faroes basalt, 2002–2005) is funded by the sindri group as part of the programmes for licensees within the faroese offshore area, and addresses these issues with special focus on the subaerially extruded flood basalts of the faroe islands (cf. japsen et al. in press). seismic and petrophysical properties of faroe islands basalts: the seifaba project peter japsen, morten sparre andersen, lars ole boldreel, regin waagstein, robert s. white and michael worthington fig. 1. a: geological map of the faroe islands showing the location of deep boreholes and the distribution of the three palaeogene basalt formations (modified from waagstein 1998). b: location of the faroe islands relative to the extent of flood basalts and producing oil fields (modified from grant et al. 1999; sørensen 2003). geological survey of denmark and greenland bulletin 4, 53–65 (2004) © geus, 2004 drilling and logging at glyvursnes and vestmanna the glyvursnes-1 well was drilled outside tórshavn to 700 m as a slim borehole with wire-line coring technique as part of the seifaba project in 2002 (fig. 1; waagstein et al. 2003). glyvursnes is well suited for combining vertical seismic profiles (vsp) and surface seismic experiments (onshore and offshore): the terrain is relatively flat and the seismic effects of a nearby near-vertical shear zone can be studied in detail (fig. 2). during the same operations, the existing 660 m deep vestmanna-1 well was reamed and logged. three deep wells on the faroe islands have drilled thick sections of the lower, middle and upper basalt formations (fig. 3; rasmussen & noe-nygaard 1970): 1. the vestmanna-1 well was drilled to 660 m in 1980 and reopened to 590 m in 2002 and logged. the hole was drilled in the lower part of the middle basalt formation and extends 100 m into the lower basalt formation. a full core was taken. 2. the lopra-1/1a well was drilled through the lower basalt formation to 2.2 km in 1981, and deepened to 3.6 km in 1996 without reaching the base of the volcanic succession. one short core of basalt was taken from this well. 3. the glyvursnes-1 well was drilled to 700 m in 2002 to around the boundary between the middle and upper basalt formations (fig. 4). a full core was taken. an extensive logging programme was run in the glyvursnes and vestmanna boreholes, comparable to that previously run in the lopra well in 1981 and 1996. the programme includes optical televiewer, caliper, natural gamma ray, resistivity, neutron porosity, density, full wave sonic, spectral gamma (of poor quality) and temperature/conductivity measurements. the vestmanna and glyvursnes boreholes penetrate, respectively, the lowermost 550 m and the uppermost 450 m of the 1400 m thick middle basalt formation. in both sections the middle basalt formation is characterised by the presence of plagioclase-phyric compound flows composed of thin flow-units of variable porosity, and by rare thin beds of tuff. the lowermost part of the middle basalt formation in the vestmanna borehole and the lowermost 70 m of the upper basalt formation in the glyvursnes borehole are very similar. the overlying section at about 230 to 285 m in glyvursnes-1 consists of a few relatively thick plagioclasephyric flow-units possibly forming two compound flows (the tórshavn flows), that are morphologically very similar to the flows of the lower basalt formation (except that the latter are near-aphyric). future work ultrasonic velocities and other parameters will be measured on core samples from selected drill holes. the samples will be investigated under varying pressure for both dry and saturated samples. studies of the velocity–porosity relationships of basalts will focus on matrix properties, but will also take into account variations in magma type, secondary mineralisation, pore shapes and fractures. a lithostratigraphic interpretation of the borehole logs will be carried out, and the results will be correlated with petrography, rock chemistry and ultrasonic properties of core samples. this comparison of data acquired at core and log scale will be extended to data of seismic scale acquired from flood basalts of all three formations. well log and core data 54 fig. 2. aerial photograph of glyvursnes with well location and outline of seismic data recording on glyvursnes in 2003. red crosses indicate the locations of permanent seismometers during the experiments. blue dots mark the positions of the three airgun pits for the vsp experiments. blue lines mark the locations of the 14 hz geophone strings used for the onshore reflection profiles. red lines show the locations of onshore geophone strings and the tethered streamer during the offshore–onshore experiments and marine reflection experiment. yellow and green crosses mark shotpoints for the marine reflection experiment (from japsen et al. in press). will be examined to evaluate how magma type and secondary mineralisation influence the relationships between velocity and porosity of basalt. supplementary analysis of well logs from exploration wells in the faroe–shetland basin will be carried out to provide additional data concerning the distribution of elastic properties of the basalts around the faroe islands. seismic experiments at glyvursnes the extrusive basalt flows in the faroe–shetland basin are generally sub-horizontal on a regional scale, although there are often large physical property variations (largely governed by porosity) within individual flows, and locally rugged small-scale relief on the tops of flows. the strong layering and local relief may cause internal multiples, forwardand backscattering of the incident energy, multiple-mode conversion, anisotropy, absorption and geometric spreading, and lowpass filtering of the energy that propagates through a stacked layer of basalt flows. since flow thicknesses are at least an order of magnitude smaller than the seismic wavelength, reflections are rarely seen from individual flows: the seismic response depends on the complex interactions of reflections from multiple flow units. to study these effects, three closely integrated seismic experiments were carried out around the glyvursnes-1 borehole in 2003 (fig. 2). three vsp surveys were undertaken with source offsets of 14, 242 and 415 m. the two offset locations were on either side of a well-defined shear zone. onshore and offshore high resolution data were acquired along the lines shown in fig. 2. two 625 m lines were shot with the charges along the line, while short profiles were acquired with the source displaced laterally. the marine seismic reflection data were acquired on a tethered streamer using a moving airgun cluster. additional data were obtained at longer offsets. a dense array of 45 autonomous seismometers was deployed for a six-month period around the site of the borehole. during periods of controlled source seismic shooting, all sites maintained a sampling rate of 200 samples/sec. for intervening periods, when recording earthquakes, a sampling rate of 100 samples/sec was used. in addition, three 400 m long independent temporary land arrays were set up in september 2003 using a mixture of oneand three-component geophones. throughout acquisition of the marine reflection seismic and the offshore–onshore, wide-angle seismic data, a threecomponent borehole seismometer was held clamped at a depth of 400 m in the middle basalt formation in the glyvursnes borehole. 55 fig. 3. stratigraphical position of deep boreholes in the faroe islands (modified from waagstein 1988). see fig. 1 for location of wells. fig. 4. composite log from the glyvursnes-1 well. note the correlation between p-slowness (p-wave velocity) and density/neutron porosity that reflects the porosity-contrast between the porous crust and the massive core of the basalt flows (from japsen et al. in press). future work a detailed analysis of seismic wave propagation through a typical faroese basalt succession will be carried out using the logging carried out in 2002, ongoing laboratory work on the core and the seismic data acquired during the summer of 2003. with the sonic logging and the vertical vsp, a detailed vertical velocity profile will be obtained in the glyvursnes-1 borehole and determine intrinsic and extrinsic attenuation/scattering of seismic waves in one dimension. by combining the two detailed fixed offset vsps and variable offset with data registered at fixed depth, it is planned to address attenuation/scattering in two (three) dimensions. using the multichannel seismic data and the autonomous seismic array it is planned to obtain two-dimensional and three-dimensional velocity distribution in a small area around the borehole. the multichannel surface seismic experiment is designed to analyse sub-critical and post-critical reflections from the basalt succession drilled at glyvursnes, providing an image of structural continuity in the area. perspectives drilling of the new borehole at glyvursnes and re-logging of vestmanna-1 in combination with the extensive data set for the lopra-1 well will give valuable new stratigraphic and petrophysical control of the lower, middle and upper basalt formations on the faroe islands. the planned experiments will provide a link from hand-specimen scale, through the slightly larger averaging of borehole logs, to seismic scales ranging from high-resolution vsps and multi-channel land arrays, to onshore–offshore shooting using wide-band airguns, to the truly long-period response of the basalt flows using teleseismic arrivals. the investigations will provide a unique data set and, hopefully, new understanding of the seismic and petrophysical properties of faroe islands basalts. acknowledgements thanks are due to the sindri group for permission to publish results from the glyvursnes-1 well. seifaba is funded collectively by all oil companies operating in the faroe islands sector (the sindri group). references boldreel, l.o. & andersen, m.s. 1993: late paleocene to miocene compression in the faroe–rockall area. in: parker, j.r. (ed.): petroleum geology of northwest europe: proceedings of the 4th conference, 1025–1034. london: geological society. grant, n., bouma, a. & mcintyre, a. 1999: the turonian play in the faroe-shetland basin. in: fleet, a.j. & boldy, s.a.r. (eds): petroleum geology of northwest europe: proceedings of the 5th conference, 661–673. london: geological society. japsen, p. et al. in press: preliminary results of petrophysical and seismic properties of faroes basalts (seifaba project). in: doré, a.g. & vining, b. (eds): petroleum geology: north-west europe and global perspectives: proceedings of the 6th petroleum geology conference. london: geological society. rasmussen, j. & noe-nygaard, a. 1970: geology of the faroe islands. danmarks geologiske undersøgelse, i. række 25, 142 pp. sørensen, a.b. 2003: cenozoic basin development and stratigraphy of the faroes area. petroleum geoscience 9, 189–207. waagstein, r. 1988: structure, composition and age of the faeroe basalt plateau. in: morton, a.c. & parson, l.m. (eds): early tertiary volcanism and the opening of the ne atlantic. geological society special publication (london) 39, 225–238. waagstein, r. 1998: a geological field guide to the palaeogene flood basalts of su∂eroy, faroe islands. danmarks og grønlands geologiske undersøgelse rapport 1998/30, 46 pp. waagstein, r., boldreel, l.o. & andersen, c. 2003: an integrated petrophysical approach to the sub-basalt imaging problem using well logging data to link measurements from cores and seismic surface experiments. geophysical research abstracts 5(09388), 2 pp. white, r.s., smallwood, j.r., fliedner, m.m., boslaugh, b., maresh, j. & fruehn, j. 2003: imaging and regional distribution of basalt flows in the faroe-shetland basin. geophysical prospecting 51, 215–231. authors’ addresses p.j. & r.w., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k. e-mail: pj@geus.dk m.s.a., university of the faroe islands, noatún 3, fo-100 tórshavn, faroe islands. present address: geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k. l.o.b., geological institute, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. r.s.w., bullard laboratories, university of cambridge, madingly road, cambridge cb3 0ez, uk. m.w., department of earth sciences, university of oxford, parks road, oxford ox1 3pr, uk. 56 << /ascii85encodepages false /allowtransparency false /autopositionepsfiles true /autorotatepages /all /binding /left /calgrayprofile (dot gain 20%) /calrgbprofile (srgb iec61966-2.1) /calcmykprofile (u.s. web coated \050swop\051 v2) /srgbprofile (srgb iec61966-2.1) /cannotembedfontpolicy /warning /compatibilitylevel 1.4 /compressobjects /tags /compresspages true /convertimagestoindexed true /passthroughjpegimages true /createjdffile false /createjobticket false /defaultrenderingintent /default /detectblends true /colorconversionstrategy /leavecolorunchanged /dothumbnails false /embedallfonts true 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<feff4f7f752890194e9b8a2d5b9a5efa7acb76840020005000440046002065874ef65305542b8f039ad876845f7150cf89e367905ea6ff0c4fbf65bc63d066075217537054c18cea3002005000440046002065874ef653ef4ee54f7f75280020004100630072006f0062006100740020548c002000520065006100640065007200200035002e0030002053ca66f465b07248672c4f86958b555f3002> >> >> setdistillerparams << /hwresolution [2400 2400] /pagesize [595.000 842.000] >> setpagedevice geological survey of denmark and greenland bulletin 26, 2012, 85-88 85 rock phosphate and lime for small-scale farming in tanzania, east africa per kalvig, niels fold, jesper bosse jønsson and elisante elisaimon mshiu poor soils are a major cause of poverty in sub-saharan africa, and thus restoration of soil fertility is a significant challenge for sustainable agriculture. some of the main resources required, e.g. phosphate and lime, are present in many african countries and can be used by smallholder farmers in a relatively unprocessed form instead of expensive commercial fertilisers. here we present a small study of the mbeya region in tanzania, which locally has both phosphate and lime. most soils in sub-saharan africa are losing nutrients necessary for sustainable agriculture. this is mainly due to intensive farming and the fact that the nutrients are not replaced adequately. further reasons for nutrient losses are leaching, soil erosion and fixation by iron and aluminium oxides. vast areas experience moderate to acute phosphorus deficiency (vanlauwe & giller 2006). the mbeya region in south-western tanzania (fig. 1) is characterised by intensive smallholder plots along with several local sources of phosphate-bearing rocks and limestone. the former were examined in the 1980s (chesworth et al. 1988, 1989), but have never been utilised (kalvig et al. 2010). phosphates and lime – opportunities and constraints soils may become acid for many reasons and high rainfall may lead to washing out the nutrients needed for healthy plant growth. thus phosphorus and calcium deficiencies are common factors that restrict plant growth in highly leached tropical soils. the majority of smallholders cannot afford to use adequate quantities of commercial fertilisers and lime, resulting in low yields. an alternative is to use local sources, which can improve agricultural productivity by slowly releasing essential elements and raising the ph value of the soil. phosphates, lime, potassium-rich minerals, clay, zeolite and mica are common locally, but lacking awareness of their effects, very few smallholders use them. if such an awareness could be fostered and local resources made available at affordable prices, it would give farmers an opportunity to improve their crop yields (van straaten 2002; mitchell 2005). the phosphate potential in tanzania was outlined twenty years ago by mchihiyo (1991). the study presented here shows that local phosphate can be made available at affordable costs. phosphate minerals. natural fertilisers are available and can be used untreated. they comprise organic fertilisers such as manure, leaf litter and sludge, and rock fertilisers such as marl, rock phosphate, volcanic rock and mica. in contrast to natural fertilisers, artificial fertilisers are readily soluble and contain guaranteed total, active nutrient concentrations. the practical challenges and potential benefits involved in the use of rock phosphate have been widely discussed (mchihiyo 1991; appleton 2002; van straaten 2002, 2006; vanaluwe & giller 2006). the general view is that the use of rock phosphate for local agriculture is justified, provided its addition is managed in accordance with the type of crop and the conditions of the soil. usually, crop yields only show a slow response over 2–3 years after the addition of rock phosphates, which makes it difficult to generate interest among local farmers. the solubility of phosphate-bearing rocks differs widely depending on the mineralogy and chemistry of the rock i i i i i i i i i i i i i i i i i i i i i i i i i i rukwa rift 33°e kenya tanzania 400 km major road minor road railway regional capital district capital international boundary prospects 25 km river 9°s 33°e panda hills songwe scarp mbeya mbalizi lake rukwa njelenje sukumavera muvwa lake nyasa 9°s nanyala fig. 1. map of south-western tanzania showing the mbeya region with its main towns, roads and railway. the prospects indicated are known carbonatite occurrences which were considered potential phosphate resources by previous studies. © 2012 geus. geological survey of denmark and greenland bulletin 26, 85–88. open access: www.geus.dk/publications/bull 8686 type, and not all of them are therefore effective when applied directly to the soil. some important factors for the successful use of rock phosphate are: (1) the type of phosphatic rock, with apatite as the most common phosphate mineral containing 34–42% p2o5, (2) the reactivity of the phosphate rock, (3) the soil, where in particular ph, cation exchange capacity, phosphorus and calcium concentration, and phosphorus-fixing capacity of the soil play important roles, (4) the type of crop, because the ability to use h2po4, which dominates at low ph, differs from crop species to crop species, and (5) fertiliser management (van straaten 2002). lime. the primary reason for increasing the soil ph by liming is to reduce the aluminium toxicity to plants, and to ensure that ca, mg, k and p are available to the plant. raising the ph value to 5.2–5.5 is adequate for increasing crop yields, whereas higher values reduce the breakdown of rock phosphate (mitchell 2005). phosphate and lime resources in the mbeya region geological setting. in tanzania, a central archaean craton is surrounded by proterozoic fold belts (fig. 2). the mbeya region lies south-west of the archaean craton in the nw–sestriking 2000–1800 ma old ubendian fold belt. this fold belt is dominated by gneisses and amphibolitic rocks, intruded by 730 ma old syenite-gabbro complexes, all of which have been intruded by carbonatite complexes of cretaceous to late palaeogene age. during the late phanerozoic, rifting and faulting occurred in connection with the formation of the east african rift valley. the occurrence of neogene sediments such as conglomerate, clay and chert as well as volcanic rocks is related to the rift valley. phosphate in the mbeya region. the most promising phosphate occurrence is the songwe scarp carbonatite, which is a 20 km long, 50 m wide, nw–se-striking, 100 ± 10 ma old ferrocarbonatite (miller & brown 1963; brown 1964). mchihiyo et al. (1992) provide an overview of the exploration history of this carbonatite, which took place in the 1950s. the exploration focussed on uranium and comprised geological mapping of anomalous zones of the carbonatite some 2 km north-west of the village njelenje (fig. 1; brown 1964). in the 1990s, the phosphate was assessed as a potential source of fertiliser (van straaten 2002). the exploration led to the recognition that several parts of the songwe scarp carbonatite are relatively rich in apatite. it was also discovered that areas with elevated concentrations of radioactive elements, yttrium and other rare-earth elements occur near njelenje (fig. 3). the highest phosphate concentrations were found in limonitic residual soils overlying the carbonatite, with up to 20% p2o5 and 6% k2o. these soils can be used as phosphate fertilisers (chesworth et al. 1989; appletorn 2002; van straten 2002). in addition to the songwe scarp carbonatite, the region holds several other occurrences of phosphate. one is the panda hills carbonatite where c. 1 mt of residual phosphates with a p2o5 concentration up to 10.3% have been mapped (van straaten 2002). another is the mbalizi carbonatite where weathered parts locally form a 0.5 m thick limonitic crust with up to 30% p2o5 (mchihiyo 1991; mchihiyo et al. 1992). this phosphate is under exploration as a potential source of niobium and other rare-earth elements. a third possibility of phosphate is guano from bats around sukumavera, but the amount is far too small to play any role, even for smallholders (van straaten 2002). limestone in the mbeya region. calcareous sedimentary rocks, mainly travertine in the songwe valley, are mined for (1) dimension stones using a yellowish travertine near nanyala in the mbozi district, (2) carbonate for calcination and (3) cement production. hochstein et al. (2000) and roberts et al. (2004) provided geological details of this limestone unit and estimated that it covers an area of c. 13 km2 with a thickness of 5–70 m, equivalent to >150 000 000 m3. an example of a weathered occurrence is shown in fig. 4. bukobau supergroup cenozoic cretaceous proterozoic 200 km 30°e 10°s 10°s tanzania craton mbeya sediments neogene volcanics proterozoic fold belts carbonatites archaean craton victoria lake lake kenya tanganyika zambia rwanda fold belt ubendian fig. 2. simplified geological map of tanzania (based on van straaten 2002). the mbeya region (framed) is located in south-western tanzania. 87 proposals for feasibility studies small-scale phosphate operation to supply local smallholders. the songwe valley area holds two potential sources for the production of local rock phosphate: carbonatite and apatite-rich residuals of carbonatite. given that previous studies indicate that enrichment of phosphorus has taken place in the residual apatite-rich soil (mchihiyo et al. 1992) and that a production based on such soil is technically relatively straightforward, the present study only considers this source. in the vicinity of the village of njelenje, the p2o5 concentration reaches 18–20% and the k2o concentration 6% (mchihiyo et al. 1992). a pilot study was conducted in march 2010. the purpose was to learn from smallholding farmers what their local cultivation practices are and get their views on requirements on and limits to the use of locally produced phosphate fertilisers.  the study also aimed at clarifying the availability of local labour for small-scale phosphate extraction. the village of muvwa, located in the mbeya region, was chosen due to its proximity to the phosphate resources of the songwe scarp carbonatite (fig. 3). the study area encompassed 420 households corresponding to 1768 inhabitants of whom 331 persons were capable of working. twenty-one households were interviewed in order to learn about the general pattern of cultivation practices. the low number means that the results are only indicative (kalvig et al. 2010). as no geological data on the survey area are available, the data from njelenje, situated c. 5 km to the north of the interview area, may serve to indicate the total available phosphorus content of the soils in the survey area. a feasibility study of a potential rock phosphate production is warranted and, based on a labour intensive concept producing rock fertiliser for local consumption, should (1) assess the phosphate resource (grade and tonnage estimates – and the potential content of harmful elements), (2) propose a suitable set-up for the production and (3) assess the distribution and market prospects. at this stage, no calculations of the required amount of local rock phosphate per hectare can be made, because a number of technical data are not yet available: the actual fertility of the soils of the trial area; depletion rate of phosphorus; fixation rate of phosphorus; identification of the most phosfig. 3. landscape around the village of njelenje in the songwe valley, looking east and showing the songwe scarp carbonatite. fig. 4. small section of limestone in the nanyala area of the mbeya region showing the boundary between topsoil and weathered rock. 8888 phorus demanding crops; the average phosphorus content of the potential rock phosphate. small-scale extraction of lime to supply local smallholders. in the village of nanyala, some hundred people are involved in small-scale mining of a weathered, whitish travertine accompanied by a semi-mechanised production of hydrated lime. more than ten mining licences cover the lime producing area. the limestone resource in the songwe valley consists of various grades of quaternary to recent travertine deposits. it is dominated by poorly consolidated, partly weathered, whitish rock, of which neither the quantity nor the quality has ever been thoroughly investigated as regards lime for agricultural use. the ideal agricultural lime is a ground dolomite or dolomitic limestone with a particle size <2 mm; 60% <400 μm and up to 50% <150 μm (mitchell et al. 1997). labour intensive production of agricultural lime is a relatively straightforward process (mitchell & mwanza 2005). the march 2010 survey included interviews with some of the operators in the neighbouring nanyala area (kalvig et al. 2010). hydrolime has been produced for several years in this area. there is a potential for diversifying the existing hydrolime production to include lime products for local use and particularly for sale to local smallholders cultivating coffee. in tanzania, agricultural lime is mainly produced in the tanga district located c. 800 km from the mbeya region. a feasibility study on how to extend the on-going production of hydrolime to include lime for agricultural use seems warranted. it should include all steps from quarrying and marketing to testing the products and should be based on a labour intensive concept. acknowledgements geocenter denmark is thanked for financial support, and the tanzania commission for science and technology for permission to conduct field work. references appleton, j.j. 2002: local phosphate resources for sustainable development in sub-saharan africa. british geological survey report cr/02/121/n, 134 pp. brown, p.e. 1964: the songwe scarp carbonatite and associated feldspathization in the mbeya range, tanganyika. quarterly journal of the geological society 120, 223–240. chesworth, w., semoka, j.m.r., van straaten p., mnkeni, p.m.s., kamasho, j.a.m. & mchihiyo, e.p. 1988: tanzania–canada agrogeology project: report on completion of the first phase, 93 pp. ontario, canada: university of guelph. chesworth, w., van straaten, p. & semoka, j.m.r. 1989: agrogeology in east africa: the tanzania–canada project. journal of african earth sciences 9, 357–362. hochstein, m.p., temu, e.p. & moshy c.m.a. 2000: geothermal resources of tanzania. world geothermal congress 2000, kyushu–tohoku, japan, 28 may to 10 june 2000. proceedings 1233–1238. kalvig, p., fold, n., jønsson, j.b. & mshiu, e.e. 2010: local use of agrominerals. untapped resources for farming communities in sub-saharan africa. appraisal study on the agromineral potential in the mbeya area, tanzania. danmarks og grønlands geologiske undersøgelse rapport 2010/107, 58 pp. mchihiyo, e.p. 1991: phosphate potential in tanzania. fertilizer research 30, 177–180. mchihiyo, e.p., kiranga, o.w. & mbasha, m.z. 1992: tanzania –canada agromineral project. report on completion of second phase (madini group). unpublished report, madini group, tanzania. miller, j.a. & brown, p.e. 1963: the age of some carbonatite activity in south-west tanganyika. geological magazine 100, 276–279. mitchell, c.j. 2005: farm lime: low cost lime for small-scale farming. british geological survey technical report cr/03/066n, 138 pp. mitchell, c.j. & mwanza, m. 2005: manual for small-scale production of agricultural lime. british geological survey commissioned report cr/05/092n, 31 pp. mitchell, c.j., inglethorpe, s.d.j., tawodzera, p., bradwell, s. & evans, e.j.m. 1997: local development of affordable lime in southern africa. british geological survey technical report wc/94/20, 120 pp. roberts, e.m., o’connor, p.m., gottfried, m.d., stevens, n., kapalima, s. & ngasal, s. 2004: revised stratigraphy and age of the red sandstone group in the rukwa basin, tanzania. cretaceous research 25, 749–759. van straaten, p. 2002: rocks for crops. agrominerals of sub-saharan africa, 338 pp. nairobi, kenya: international centre for research in agroforestry. van straaten, p. 2006: farming with rocks and minerals: challenges and opportunities. anais da academia brasileira de ciências 78, 731–747. vanaluwe, b. & giller, k.e. 2006: popular myths around soil fertility management in sub-saharan africa. agriculture, ecosystems and environment 116, 34–46. authors’ addresses p.k., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: pka@geus.dk n.f., department of geography and geology, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. j.b.j., school of geographical & earth sciences, university of glasgow, glasgow g12 8q q, scotland, uk. e.e.m., department of geology, university of dar es salaam, p.o. box 35052, dar es salaam, tanzania. geological survey of denmark and greenland bulletin 28, 2013, 25-28 25 terrain subsidence detected by satellite radar scanning of the copenhagen area, denmark, and its relation to the tectonic framework peter roll jakobsen, urs wegmuller, ren capes and stig a. schack pedersen in the european union (eu) project terrafirma, which is supported by the european space agency to stimulate the global monitoring environment system, we are using the latest technology to measure terrain motion on the basis of satellite radar data. the technique we employ is known as persistent scatterer interferometry (psi); in denmark, it was previously used to map areas of subsidence susceptible to flooding in the danish part of the wadden sea (vadehavet) area (pedersen et al. 2011). that study was part of the flooding risk theme under the terrafirma extension project. another coastal protection monitoring activity in the eu seventh framework project subcoast followed, in which the low-lying south coast of lolland, prone to flooding, was studied. the geological survey of denmark and greenland (geus) is also involved in the three-year eu collaborative project pangeo in which geus is one of 27 eu national geological surveys. the objective of pangeo is to provide free and open access to geohazard information in support of the global monitoring environment system. this will be achieved by providing a free, online geohazard information service for the two largest cities in each eu country, i.e. 52 towns throughout europe with c. 13% of eu’s population. the danish cities selected for investigation under terrafirma are copenhagen and aalborg. capitals have first priority, and aalborg was chosen because of good satellite data. in this paper, psi data for copenhagen are presented together with interpretations of terrain displacement (fig. 1). psi processing of satellite radar data for copenhagen the satellite data covering copenhagen were obtained from the descending track d480 ers satellite in the period 1992–2000. the psi processing was carried out by gamma remote sensing ag, using a method that was carefully qualified and validated in the terrafirma project (crosetto et al. 2008). geus carried out the analysis using the program arcgis, in which geological and topographical data provide the basis for interpretation. based on the psi data, nine areas were outlined in which subsidence had occurred over the period 1992–2000. the areas are between 0.1 and 2.2 km2 and here we present two of them. apart from those mentioned above, minor subsidence differences of regional extent have been recognised; these are interpreted as tectonic. the copenhagen area that was processed has a size of 960 km2 with a reference point at 55.685668°n, 12.493937°e. a total of 419  660 psi points were identified, corresponding to 437 points per km2. the majority of the points (94.5%) show small rates of vertical motion, i.e. between –1.5 and +1.5 mm/year. a small number of points (1.6%) show subsidence rates of 3.5 to 1.5 mm/year, and a few (0.2%) show subsidence rates of more than 3.5 mm/ year. a few points (0.4%) show uplift rates between 1.5 and 3.5 mm/year; these are regarded as scattered uncertainties in relation to the average annual motion rate of 0.35 mm/ fig. 1. map of greater copenhagen showing the area covered by psi data. vertical movements are represented in a raster grid showing the average movement in 100 × 100 m squares. the concentrations of yellow-red colours show areas with maximum rates of subsidence. note the regional difference in light and dark green colours which might be caused by tectonic subsidence east of the carlsberg fault zone. © 2013 geus. geological survey of denmark and greenland bulletin 28, 25–28. open access: www.geus.dk/publications/bull carlsberg fault zone rate of change (mm/year) −10 to −4 −4 to −3 −3 to −2 −2 to −1 −1 to −0.5 −0.5 to 0.5 0.5 to 1 1 to 2 2 to 3 3 to 4 4 to 5 5 km 55°40´n 12°28´e øresund sjælland amager furesø airport harbour 2626 year for the entire area, with a standard deviation of annual motion rate amounting to 0.74 mm/year. the geology of copenhagen and its relation to subsidence areas copenhagen is located on the east coast of the island of sjælland, which is separated from sweden by the strait of øresund (fig. 1). part of the city extends onto the smaller, neighbouring island of amager, and the strait between the two islands is the site of copenhagen harbour. the airport of copenhagen, kastrup, is situated on the southern part of amager. most parts of copenhagen are lowland, i.e. a few metres above sea level, but the terrain rises northwards and westwards where it reaches heights of 50 m a.s.l. the bedrock of copenhagen is dominated by danian limestone. two units are found: the stevns klint formation that comprises bryozoan limestone rich in chert beds (surlyk et al. 2006) and the københavn kalk formation which is dominated by calcarenitic, calcilutitic limestone with chert beds (stenestad 1976). an important tectonic feature is the carlsberg fault zone (stenestad 1976; jakobsen et al. 2002) that can be followed from the south coast of amager northwestwards to furesø (figs 1, 2). the quaternary deposits of copenhagen comprise ice age and postglacial deposits. the latter consist of terrestrial sediments that accumulated in streams and bogs, and marine sand and gravel which accumulated along the coasts (fig. 2b). the ice-age deposits are dominated by a widespread young till unit overlying meltwater sand and gravel, and more local, older till units and meltwater clay. tunnel valleys, depressions in hummocky moraine and stream valleys form wetlands around copenhagen, where freshwater deposits, mainly peat, accumulated in the holocene. the western part of amager is reclaimed sea floor with marine and coastal deposits. areas with dump and fill deposits occur along the coasts or in peat-dominated depressions and may be characterised by high rates of subsidence. examples of areas with subsidence based on the geology and records of man-made ground, there are three types of areas with potential risk of subsidence, namely areas underlain by postglacial peat deposits (amounting to 70 km2), large areas of man-made ground (53 km2) and small areas of reclaimed land (19.4 km2). one of the areas with subsidence identified from the psi data in the period 1992–2000 is lersøparken (figs 2, 3) with fig. 2. correlation between ground stability and surficial deposits in greater copenhagen. a: map of classified soft ground areas in greater copenhagen. the rectangles show the location of two examples of subsiding areas described in the text. b: geological map of surficial deposits. the fill and reclamation areas did not exist when the region was mapped in 1899 (rørdam 1899). ]] amager a b furesø 10 km fig. 3 fig. 4 55°40´n 12°30´e observed psi, natural ground movement observed psi, anthropogenic artificial ground potential instability, natural ground potential instability, anthropogenic artificial ground clayey till meltwater sand meltwater gravel aeolean sand peat freshwater sand freshwater clay unmapped marine sand marine gravel marine clay fill and land reclamation 27 a subsidence rate between 2 and 5 mm/year. the 0.22 km2 subsiding area is situated in a ne–sw-trending valley where postglacial peat has accumulated above clay in an elongated depression (fig. 3). the depression was used as a dump site in the period from 1880 to 1920. peat is easily compressed, and compaction of the waste is an additional factor accounting for the high subsidence rate. kalveboderne with valbyparken and tippen along the coast of southern copenhagen is another example of an area with subsidence (figs 2, 4). a considerable number of psi points show subsidence of more than 4 mm/year. this area of 1.4 km2 was used as a dump from 1913 to 1960 when waste was dumped on the beach and in the adjacent shallow sea. in 1961, the area was extended with fill to the present artificial shoreline. compaction of the soft, natural sediments and waste followed by fill deposits lead to subsidence. subsidence related to tectonic features the most important tectonic feature in the subsurface of copenhagen is the se–nw-striking carlsberg fault zone (fig. 1). the fault is part of a number of relay faults related fig. 3. detailed map of lersøparken (red frame). a: orthophotograph of the area with psi points representing areas of subsidence. b: geological map of the area showingt clay and peat in the ne–sw-striking valley. for location see fig. 2. fig. 4. detailed map of the kalveboderne area where waste and fill have been deposited on holocene marine and coastal deposits. a: orthophotograph with psi data points. b: an old topographical map showing the same area prior to deposition of waste and fill. for location see fig. 2. ! ! !! ! !!!!! ! !! !!!!!! ! ! ! ! !!! !!! ! ! ! !!! ! !!! !!! ! ! ! ! !! !! ! ! !! !!! ! !!! !!!!! velocity clayey till peat freshwater clay500 m > –5 –4 to –5 –3 to –4 –2 to –3 –2 to –0.75 0.75 to –0.75 0.75 to 2 55°43´n 12°34´erate of change (mm/year) aa bb 55°37´37´´n 12°31´e tippen 500 m > –5 –4 to –5 –3 to –4 –2 to –3 –2 to –0.75 0.75 to –0.75 0.75 to 2 a b valbyparken rate of change (mm/year) 2828 to the tornquist–sorgenfrei wrench fault zone. a seismic cross-section of the carlsberg fault zone shows that it can be classified as a negative flower structure with a mean vertical offset of 50–100 m of the limestone deposits. the hanging block is found north-east of the fault zone (fig. 5; fallesen 1995; jakobsen et al. 2002). the limestone in the fault zone itself is strongly fractured as documented by low seismic velocity in the fault-affected zone (nielsen et al. 2005). there is clear evidence of weak, regional subsidence east of the fault zone, i.e. the area of the down-thrown fault block (fig. 1). the carlsberg fault zone can be followed north-westwards to furesø, which is the deepest lake in denmark, and we suggest that the shape of the lake is governed by displacement along the fault zone. this would be an alternative explanation of the origin of the lake, which has hitherto been regarded as formed from a combination of tunnel valleys and kettle holes. the subsidence recorded by the psi points may correspond to subsidence rates in the copenhagen area recorded from traditional levelling (mark & jensen 1982). groundwater extraction may also influence subsidence, which could have been the case for amager. however, the groundwater level on amager was stable during the period of the satellite data acquisition. references crosetto, m., monserrat, o., adam, n., parizzi, a., bremmer, c., dortland, s., hanssen, r.f. & van leijen, f.j. 2008: validation of existing processing chains in terrafirma stage 2, final report, 15 pp., http://www. terrafirma.eu.com/validation/valproj/final%20reports/valproj _final_report.pdf. fallesen, j. 1995: stratigraphy and structure of the danian limestone on amager, examined with geophysical investigations – especially with regard to the carlsberg fault. unpublished msc thesis, university of copenhagen, denmark jakobsen, p.r., fallesen, j. & knudsen, c. 2002: strukturer i den københavnske undergrund – folder, forkastninger og sprækker. dansk geoteknisk forening bulletin 19, 19–29. mark, a. & jensen, j.e. 1982: niveauændringer i københavn bestemt ud fra gentagne præcisionsnivellementer. landinspektøren 32, 10–21. nielsen, l., thybo, h. & jørgensen, m.i. 2005: integrated seismic interpretation of the carlsberg fault zone, copenhagen, denmark. geophysical journal international 162, 461–478. pedersen, s.a.s., cooksley, g., gaset, m. & jakobsen, p.r. 2011: detection of terrain changes in southern denmark using persistent scatterer interferometry. geological survey of denmark and greenland bulletin 23, 41–44. rørdam, k. 1899: beskrivelse til geologisk kort over danmark, (i maalestok 1:100,000). kortbladene kjøbenhavn og roskilde. danmarks geologiske undersøgelse i. række 6, 88 pp. stenestad, e. 1976: københavnsområdets geologi især baseret på citybaneundersøgelserne. danmarks geologiske undersøgelse iii. række 45, 149 pp. (with summary in english). surlyk, f., damholt, t. & bjerager, m. 2006: stevns klint: uppermost maastrichtian chalk, cretaceous–tertiary boundary, and lower danian bryozoan mound complex. bulletin of the geological society of denmark 54, 1–48. top cretaceous 100 m sw ne top cretaceous 0 100 200 300 tw ow ay tr av el ti m e (m se c) fig. 5. seismic cross-section of the carlsberg fault zone (from jakobsen et al. 2002). maastrichtian chalk is found at the top of the western block, whereas danian limestone is present at the top of the eastern block. authors’ addresses p.r.j. & s.a.s.p, geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: prj@geus.dk u.w., gamma remote sensing, worbstrasse 225, ch-3073 gümligen, switzerland. r.c., npa satellite mapping, crockham park, edenbridge, kent tn8 6sr, uk. http://www.terrafirma.eu.com/validation/valproj/final reports/valproj_final_report.pdf http://www.terrafirma.eu.com/validation/valproj/final reports/valproj_final_report.pdf http://www.terrafirma.eu.com/validation/valproj/final reports/valproj_final_report.pdf e2019430205-01 spaceborne remote sensing is a suitable tool for early mineral exploration and surveying large areas of high arctic environment in a fast and cost-effective manner. while spaceborne data have been used widely to map geology in arid areas, similar approaches for remotely-sensed geological mapping of arctic environments is yet to be developed. freely available spaceborne optical data provides detailed information of high-quality that could potentially reduce resource exploration risk in remote regions. to this end, this study compares the use of two different multispectral spaceborne datasets (i.e. the advanced spaceborne thermal emission and reflection radiometer (aster) and sentinel-2) to map geological units in and around wollaston forland, northeast greenland – an area rich in jurassic and cretaceous sedimentary rocks and important targets for offshore petroleum exploration. multispectral image sensors simultaneously capture image data within multiple wavelength ranges (bands) across the electromagnetic spectrum. each band is commonly described by the band number and the band wavelength centre position. here, we identify the bands most suitable for geological mapping in an arctic setting, using the wollaston forland area as an example. we compare the results obtained by processing spaceborne data with a published geological map for the area (henriksen 2003). geological setting north-east greenland comprises three main geological units – a palaeoproterozoic crystalline gneiss-granite basement, a mesoproterozoic high-grade metasedimentary unit and a slightly deformed and metamorphosed neoproterozoic – lower palaeozoic sedimentary sequence – as well as a mesozoic sedimentary and volcanic basin (fig. 1). the basal part of this sequence constitutes the eleonore bay supergroup, marine siliciclastic and carbonate sediments intruded by caledonian granites (sønderholm & tirsgaard 1993). wollaston forland is dominated by jurassic and cretaceous sedimentary rocks. the upper permian siliciclastic and carbonate sediments are overlain by shallowand deepwater sandstones and shales deposited during continuous rifting in the triassic, jurassic and cretaceous (fig. 1). the extensive paleogene plateau basalts on wollaston forland, as well as basaltic sills and dykes in the mesozoic sediments, belong to the east greenland tertiary volcanic province (henriksen et al. 2009). the upper permian to early eocene deposits and break-up related flood basalts, dikes and sills are moderately exposed, where limited exposure is, in most places, due to the erosional and depositional processes of the quaternary glaciations. spaceborne datasets the aster sensor covers a broad spectral region with 14 bands ranging from visible-near infrared (vnir) and shortwave infrared (swir) to thermal infrared (tir), with 15 m, 30 m, and 90 m spatial resolution, respectively (see fig. s1 in supplementary information). this sensor has been widely used for geological mapping applications since its launch in 1999 (abdeen et al. 2001; amer et al. 2010). sentinel-2 covers 13 bands in the vnir and swir wavelength regions, with four bands at 10 m, six bands at 20 m and three bands at 60 m spatial resolution (fig. s1). the potential for using sentinel-2 for geological remote sensing in arid and semi-arid regions has been evaluated before (al-nahmi et al. 2017; mielke et al. 2014; van der meer et al. 2014; van der werff & van der meer 2016). several pre-processing steps were applied on both datasets prior to performing the mapping approaches. the datasets were first atmospherically corrected. features associated with clouds, ice, snow and ocean were then masked out. removal of these unwanted features improved the visualisation of slight differences between spectrally similar minerals, and comparison of aster and sentinel-2 spaceborne datasets for geological mapping: a case study from north-east greenland sara salehi*1, christian mielke2, christian brogaard pedersen1 and simun dalsenni olsen1 geus bulletin is free to individuals and institutions in electronic form. the author(s) retain copyright over the article contents. research article | open access geus bulletin vol 43 | e2019430205 | published online: 17 july 2019 https://doi.org/10.34194/geusb-201943-02-05 https://doi.org/10.34194/geusb-201943-02-05 mineral mixtures such as clays and other hydrous phases resulting from alteration of primary rock-forming silicates. we adopted the following four approaches to identify the bands most suitable for differentiating rock types in the wollaston forland area: (1) enhanced false colour composite images using vnir and swir bands: a technique known as decorrelation stretch is used to remove the inter-band correlation found in the input pixels and enhance the colour differences representative of different lithological units (2) principal component analysis (pca) of aster and sentinel-2 reflectance data (i.e. vnir and swir bands) (3) classification based on band ratios (4) band-depth colour composites computed from three spectral bands (two continuum and one centre band). aster false colour composite images generating trueand false-colour composite images from aster and sentinel-2 band combinations or band ratio combinations can reveal important mineralogical/lithological information. here, we generate false colour composite images using vnir and swir reflectance bands (figs 2a, c) and tir radiance bands (fig. 2b) to separate the main lithological groups exposed in the study area. in addition, we apply a decorrelation stretch (wahi et al. 2013) on the false colour composite images, which enhances the discrimination between lithological units. in the aster vnir-swir false colour composite image (fig. 2a), exposed lithologies are highlighted by red, green and purple colours corresponding to the variable content of silicate and clay minerals and iron-oxides, respectively. rocks enriched in quartz, feldspar and silica are best detected using the aster tir false colour composite (fig. 2b). for example, migmatitic and siliciclastic metasediments in the geological map are highlighted in detail by green coloured pixels. the basaltic plateau lavas (predominantly silica oversaturated tholeiites) are not well identified in the sentinel-2 false colour composite image (sparse olive-green pixels in fig. 2c), unlike the aster tir false colour composite ice quaternary, undifferentiated east greenland palaeogene volcanic province basaltic sills and dykes basaltic plateau lavas (58–54 ma) sedimentary basins cretaceous jurassic triassic permian, upper carboniferious infracrustal units palaeoproterozoic crystalline complexes, mainly gneisses allochthonous caledonian thrust sheets (70°n–81°n) caledonian granites (c. 440–430 ma) neoproterozoic or caledonian granties (940–910 ma or c. 430 ma) neoproterozoic lower palaeozoic sediments and metasediments neoproterozoic marine shelf sequence. eleonore bay supergroup, upper part neoproterozoic siliciclastic shelf sequence, eleonore bay supergroup, lower part early neoproterozoic rocks migmatitic metasediments. krummedal sequence (c. 1000 ma) mesoproterozoic metasediments metasediments. krummedal supracrustal sequence (c. 1000 ma). may locally include neoproterozoic rocks wollaston forland kuhn ø young sund gael hamke bugt 19°w 20°w 20°w 21°w 75 ° 74 °4 0´ 74 °4 0´ 74 °2 0´ 74 °2 0´ 74 ° 20 km e2019430205-02 fig. 1. geological map of north-east greenland (modified from henriksen 2003). e2019430205-03 image, which clearly depicts these units in pink (fig. 2b). these units are rich in silica and feldspars and sentinel-2 does not have thermal bands to resolve these features. sentinel-2 is, however, better able to delineate clay, gossan and ferric-oxides (purple and blue pixels in fig. 2c) using the bands at 2190 nm and 900 nm and 560 nm, compared with the aster vnir-swir false colour composite image. principal component analysis we employed pca to enhance and separate spectral signatures from the background and to qualitatively identify differences in lithology for both datasets. pca of reflectance data provides a new layer of information that complements geological maps, e.g. by mapping offsets between lithological boundaries, which may resemble important structural or stratigraphic contacts. the first principal component (pc) band contains the highest variance, the second pc band contains second high variance and the last pc band contains the lowest variance and the highest noise (richards & xiuping 1999). not all sentinel-2 bands are suitable for geological applications and only 9 out of 13 are used in the pca. excluded bands comprise band 8 due to its wide bandwidth and bands with 60 m spatial resolution (see fig. s1) that are positioned in an atmospheric absorption feature as an aid in atmospheric correction (i.e. water vapour absorption bands and bands for cirrus cloud detection – bands 1, 9, 10). for the aster image, only vnir-swir bands are used in this step. we retrieved nine principal components for each dataset, of which, principal component bands with the largest amounts of data variance were employed to give the best discrimination between the various lithological units. in both scenarios, the results revealed geological features that were not previously identified in the geological map used in this study. band combinations of pc3 (red), pc5 (green) and pc4 (blue), were best able to delineate different lithological units in the aster images (fig. 3a). in sentinel-2 images, lithological units were best identified by band combinations of pc4 (red), pc9 (green) pc5 (blue) (fig. 3b). a b c fig. 2. false colour composite images. a: aster vnir-swir (red: band 6, green: band 4, blue: band 3). b: aster tir (red: band 10, green: band 14, blue: band 12). c: sentinel-2 (red: band 12, green: band 4, blue: band 8a). images are available in high resolution as supplementary data (fig. s2). a b fig. 3. false colour composite images based on principal component analyses (pca). a: aster data. red pixels: pc3. green: pc5: blue: pc4. palaeoproterozoic crystalline gneissgranite appears as bright yellow pixels and the neoproterozoic-lower palaeozoic sedimentary sequence as magenta. b: sentinel-2 data. red pixels: pc4. green: pc9. blue: pc5. the lava series are depicted by pink colours. mafic dykes and sills are visible in cyan. see fig. 1. for reference. images are available to download in high resolution as supplementary data (fig. s3). e2019430205-04 classification using band ratios during the past two decades, aster-derived band ratios have been successfully used as proxies for mapping mineral assemblages or individual mineral groups in arid regions (amer et al. 2010; gad & kusky 2006; pournamdari et al. 2014). band ratios are known to eliminate shadow and topographic effects from images, and therefore suit complex terrains such as those found in the arctic (salehi et al. 2019). however, they do not indicate the occurrence of a mineral with absolute certainty or with any idea of quantity, so ground truthing is essential. that said, every terrain is different, and band ratios that work in some areas for a particular high low a low b dc fe fig. 4. classification using aster data. a: albedo vs. absorption contrast band ratio b: ferrous iron (band5 / band3 + band1 / band2). c: gossan (band4 / band2). d: ferric-oxides (band4 / band3). e: sericite/muscovite/illite/smectite ((band5 + band7) / band6). f: normalised vegetation index ((band3 – band2) / (band3 + band2)), in which the red colour represents a higher relative abundance of the minerals. white areas show background values. images are available to download in high resolution as supplementary data (fig. s4). fig. 5. absorption feature depth colour combination. a: aster. b: sentinel-2. images are available to download in high resolution as supplementary data (fig. s5). wollaston forland kuhn ø b absorption feature depth at 865 nm absorption feature depth at 1610 nm absorption feature depth at 560 nm wollaston forland kuhn ø iron feature depth at 658 nm epidote feature depth at 2333 nm clay feature depth at 2165 nm a wollaston forland kuhn ø b absorption feature depth at 865 nm absorption feature depth at 1610 nm absorption feature depth at 560 nm wollaston forland kuhn ø iron feature depth at 658 nm epidote feature depth at 2333 nm clay feature depth at 2165 nm a e2019430205-05 mineral or assemblage may not work elsewhere. therefore, spaceborne data should be used in conjunction with geological maps, geochemistry, ground spectral references, and any other available data for the most accurate interpretation. figure 4a illustrates albedo vs. absorption contrast in the aster scene. dark pixels in this image correspond to areas where the overall albedo is low and/or minerals have shallow features. this means that aster band ratios in these regions cannot delineate (or poorly delineate) different lithological units. medium to high band ratio values for ferrous iron (green to red colours in fig. 4b) align well with migmatitic and siliciclastic metasediments in the geological map (dark green units in fig. 1). the results derived from gossan (fig. 4c) and ferric-oxide (fig. 4d) band ratios are similar. in some regions, medium to high values in these maps have quite high normalised vegetation index values (see red and green colours in figs 4c, d, which align with green colours in fig. 4f) and may be caused by vegetation coverage. high values in the sericite, muscovite, illite and smectite image (fig. 4e) correspond to the palaeoproterozoic crystalline gneiss-granite basement and migmatitic and siliciclastic metasediments in the geological map (see yellow, brown and green units in fig. 1). band-depth colour composites wavelength position, depth, width and asymmetry of an absorption band are the most essential information used in spectroscopy that can be directly linked to mineral types and their abundances, as well as to subtle changes in the chemical composition of minerals (van der meer 2004; van der meer et al. 2018; van ruitenbeek et al. 2014). this information is particularly useful in areas where field validation is sparse, and imagery contains shallow spectral absorption features. the depth of absorption features for iron, clay and epidote/ serpentine/carbonate minerals are calculated for the aster scene and visualised as a false colour composite image (fig. 5a). the rgb colour combination for the iron feature depth at c. 900 nm (green peak at 560 nm; sensitive to total chlorophyll in vegetation) in the aster image, and the feature depth at 1610 nm for sentinel-2 (fig. 5b), agrees best with the geological map (fig. 1). sentinel-2 allows us to highlight dykes and sills (dark blue and cyan pixels in fig. 5b). the lava series is also cleared identified (cyan colours in aster, and yellow in sentinel-2). quaternary rocks (enriched in clay/silicates) are more clearly visible in the aster images (dark blue pixels), and as purple and green colours in the sentinel-2 band-depth colour composite. both datasets provide detailed information of the spatial distribution of lithologies. for example, red pixels in fig. 5a indicate iron-rich regions that add complementary information to the geological map. conclusions in this study, aster outperformed sentinel-2 in discriminating between the lithological units in the wollaston forland area. the success can be attributed to the fact that aster has six swir bands whereas sentinel-2 has only two. moreover, sentinel-2 does not contain tir bands. however, sentinel-2 is better suited to mapping iron-bearing minerals since it has several bands that cover the 900 nm iron absorption feature, while aster has only one band covering this feature. image enhancement (decorrelation stretch) has also been shown to be effective in the identification and visualisation of different rock units. using absorption feature-depth colour composition yielded the best results in delineating different lithologies. it should be noted that the band ratio combination approach only uses three bands at a time, thus knowledge-based classification (e.g. argialas & goudoula 2003; harvey & fotopoulos 2016) using those band ratios is recommended as the next step to improve these classification results. acknowledgments we 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3, 33–45. how to cite salehi, s., mielke, c., brogaard pedersen, c. & dalsenni olsen, s. 2019: comparison of aster and sentinel-2 spaceborne datasets for geological mapping: a case study from north-east greenland. geological survey of denmark and greenland bulletin 43, e2019430205. https://doi. org/10.34194/geusb-201943-02-05 *corresponding author: sara salehi | e-mail: ssal@geus.dk 1 geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350, copenhagen k, denmark. 2 helmholtz center potsdam, gfz german research center for geoscience, 14473 potsdam, germany. e2019430205-06 https://doi.org/10.5194/isprs-archives-xli-b8-423-2016 https://doi.org/10.3390/rs6086790 https://doi.org/10.3390/rs6086790 https://doi.org/10.1016/j.asr.2014.04.022 https://doi.org/10.1007/978-3-662-03978-6 https://doi.org/10.1007/978-3-662-03978-6 https://doi.org/10.1016/j.jag.2003.09.001 https://doi.org/10.1016/j.jag.2017.09.008 https://doi.org/10.1016/j.rse.2014.03.022 https://doi.org/10.3390/rs8110883 https://doi.org/10.1016/j.pss.2014.06.009 https://doi.org/10.34194/geusb-201943-02-05 https://doi.org/10.34194/geusb-201943-02-05 mailto:ssal%40geus.dk?subject= e2019430102-01 general circulation models (gcms) are the main tools used to assess the impacts of climate change. due to their coarse resolution, with cells of 100 km × 100 km, gcms are dynamically downscaled using regional climate models (rcms) that better incorporate the local physical features and simulate the climate of a smaller region, e.g. a country. however, rcms tend to have systematic biases when compared with local observations, such as deviations from day-to-day measurements, and from the mean and extreme events. as a result, confidence in the model projections decreases. one way to address this is to correct the rcm output using statistical methods that relate the simulations with the observations, producing bias-corrected (bc) projections. here, we present the first assessment of a previously published method to bias-correct 21 rcm projections of daily temperature and precipitation for denmark. we assess the projected changes and sources of uncertainty. the study provides an initial assessment of the bias correction procedure applied to this set of model outputs to adjust projections of annual temperature, precipitation and potential evapotranspiration (pet). this method is expected to provide a foundation for further analysis of climate change impacts in denmark. material and methods climate models we analysed 21 rcms from the euro-cordex initiative (jacob et al. 2014) driven by gcms from the coupled model intercomparison project phase 5 (taylor et al. 2012). of these, 16 combinations are driven by the greenhouse gas concentration scenario (representative concentration pathway) rcp 8.5 and five are driven by rcp 4.5 (table 1). rcps are based on a review of existing scientific literature considering different descriptions of future socioeconomic conditions, technological development, the environment, climate and emission of greenhouse gases and aerosols (moss et al. 2010). rcp 8.5 represents a rising radiative forcing reaching 8.5 w/m2 by 2100 whereas rcp 4.5 represents a scenario of stabilised radiative forcing at 4.5 w/m2, both relative to preindustrial levels (van vuuren et al. 2011). the rcm daily outputs were remapped using the climate data operators – a collection of command line operators to analyse climate model data (schulzweida 2019) – to match the grids of the observed temperature (20 km) and precipitation (10 km) obtained from the danish meteorological institute (dmi). we remapped temperature using a bilinear interpolation and a conservative interpolation for precipitation. bias-correction precipitation and temperature data were bias-corrected using a distribution-based scaling method, whereby daily simulations were fitted to daily observations, as described by seaby et al. (2013). we used the double gamma distribution with a cut-off threshold set to the 90th percentile to bias-correct precipitation, and a normal distribution for temperature. bias correction has limitations. for example, the correction depends on the training period used to define the distribution parameters that will be used to bias-correct the simulated precipitation and temperature (lafon et al. 2013), biases associated with the driving data (maraun 2016) and any possible alterations in the signal of change in the projection (maraun 2013). bias correction also assumes stationarity in the trained parameters (chen et al. 2015). these and other limitations have been discussed in detail by maraun & widmann (2018). in our method, we used gridded observations from 1991 to 2010 as the training dataset. the parameters obtained during this training period were used to generate bc time series from 1971 to 2100. the correction method was cross-validated using a five-fold method (maraun et al. 2015), where five non-overlapping periods of equal length are defined. four periods were used to train the parameters and then the parameters were used to bias-correct the remaining climate change: sources of uncertainty in precipitation and temperature projections for denmark ernesto pasten-zapata*1, torben o. sonnenborg1, jens christian refsgaard1 geus bulletin is free to individuals and institutions in electronic form. the author(s) retain copyright over the article contents. research article | open access geus bulletin vol 43 | e2019430102 | published online: 24 june 2019 https://doi.org/10.34194/geusb-201943-01-02 https://doi.org/10.34194/geusb-201943-01-02 e2019430102-02 period. following this approach, cross-validated time series were developed for the entire period. potential evapotranspiration (pet) pet was estimated using the oudin formula (oudin et al. 2005), which uses temperature as the only climate input. the formula accurately reproduces the annual accumulated pet over denmark when compared to observations, but they are offset from the observed monthly distributions, and a correction parameter needs to be applied. here, we estimated daily pet using the climate model temperature (uncorrected and bc) as the input and applied the correction parameter. results and discussion we validated the bias correction method by comparing how well the uncorrected and bc models simulate the observed mean annual temperature and precipitation. then, we assessed the projected changes in temperature, precipitation, and pet by the end of this century for the whole ensemble and for each individual combination of gcm and rcm. we then assessed the contribution of individual sources of uncertainty in the projections. finally, we assessed the spatial distribution of the projected change for mean annual precipitation under rcp 8.5 by the end of the century along with a measure of its uncertainty. here, we assess the change in precipitation only, as its variation throughout the country is larger than that of temperature and pet. bias-corrected results mean annual temperature biases range from –1.2°c to +1.0°c in the uncorrected models and –0.1°c to +0.3°c in the bc models (data not shown). the mean annual precipitation (857 mm) biases range from –26% to +39% for the uncorrected models and between –3% and +5% for the bc simulations. even though pet is not a direct output of the climate models, we assessed the biases associated with it using uncorrected and bc temperature data as the input to the oudin formula. the biases associated with mean annual changes are for 2071–2100, relative to the 1981–2010 reference period for the uncorrected (raw) and bias-corrected (bc) simulations. gcm: general circulation model. rcm: regional climate model. rcp: representative concentration pathway. nv: natural variability. table 1. projected change in the mean annual temperature (t), precipitation (p) and potential evapotranspiration (pet) projected change by 2071–2100 compared to 1981–2010 raw bc included in the uncertainty analysis gcm rcm rcp nv gcm rcm ensemble rcp t p pet t p pet (°c) (mm) (mm) (°c) (mm) (mm) x canesm2 remo2015 r1i1p1 8.5 3.5 265 115 5.1 310 173 x ec-earth racmo 2.2 r1i1p1 8.5 3.0 51 97 3.4 126 112 ec-earth hirham5 r3i1p1 8.5 3.1 71 100 3.9 113 135 x ec-earth racmo 2.2 r12i1p1 8.5 3.2 92 100 3.7 144 124 x x ipsl-cm5a-mr rca4 r1i1p1 8.5 3.2 215 98 3.6 241 120 x miroc5 remo2015 r1i1p1 8.5 4.1 156 134 4.9 156 167 x x mpi-esm-lr remo2009 r1i1p1 8.5 2.5 108 70 3.3 133 104 x mpi-esm-lr rca4 r1i1p1 8.5 2.6 150 78 3.0 173 112 x mpi-esm-lr remo2009 r12i1p1 8.5 2.4 120 73 3.3 154 107 noresm1-m hirham5 r1i1p1 8.5 2.8 162 95 3.5 158 129 x hadgem2-es cclm 4.8.17 r1i1p1 8.5 4.3 73 140 4.6 75 150 x hadgem2-es hirham5 r1i1p1 8.5 3.8 176 121 4.7 200 159 x x hadgem2-es remo2015 r1i1p1 8.5 4.1 88 130 5.6 110 186 x x hadgem2-es racmo 2.2 r1i1p1 8.5 4.1 133 131 4.6 181 149 x x hadgem2-es rca4 r1i1p1 8.5 3.9 165 120 4.4 219 143 ec-earth hirham5 r3i1p1 4.5 1.6 50 55 2.1 70 75 x ipsl-cm5a-mr rca4 r1i1p1 4.5 2.0 86 39 2.3 106 79 x mpi-esm-lr remo2009 r1i1p1 4.5 1.2 –25 38 1.7 –10 58 mpi-esm-lr remo2009 r12i1p1 4.5 1.2 43 38 1.7 58 57 x hadgem2-es racmo 2.2 r1i1p1 4.5 2.5 112 77 2.8 147 91 ensemble mean change 8.5 3.3 133 105 4 165 135 4.5 1.7 53 50 2.1 74 72 ensemble standard deviation 8.5 0.6 56.8 22.4 0.8 56.9 27.4 4.5 0.5 51.9 16.9 0.5 58.5 14.7 e2019430102-03 pet (564 mm) range from –6% to +8% in the uncorrected models and +2% to +5% in the bc models. projected changes the bc simulations project higher temperatures and pet compared to the uncorrected simulations (fig. 1). in contrast, the uncorrected models project higher precipitation than the bc models. the change in temperature and pet by the end of the century is larger when driven by rcp 8.5 compared to rcp 4.5. the same is true for precipitation, but the difference between the two rcps is small. when driven by rcp 4.5, the mean of the uncorrected models projects an increase in temperature of 1.7°c by the end of the century, while the bc simulations project an increase of 2.1°c. under rcp 8.5, the uncorrected ensemble mean projects an increase of 3.3°c and the bc models project an increase of 4°c (table 1). ( ° c ) (m m ) (m m ) (m m ) (m m ) ( ° c ) a. temperature, rcp 4.5 b. temperature, rcp 8.5 c. precipitation, rcp 4.5 d. precipitation, rcp 8.5 e. potential evapotranspiration, rcp 4.5 f. potential evapotranspiration, rcp 8.5 fig 1. observations and uncorrected (raw) and bias-corrected (bc) projections under two rcp scenarios. mean annual temperature under a: rcp 4.5 and b: rcp 8.5. mean annual precipitation under c: rcp 4.5 and d: rcp 8.5. mean annual potential evapotranspiration under e: rcp 4.5 and f: rcp 8.5. e2019430102-04 under rcp 4.5, uncorrected models project an increase in inland precipitation of 53 mm/yr by the end of the century, in contrast to the 76 mm/yr projected by the bc models. under rcp 8.5, the uncorrected ensemble projects an increase of 133 mm/yr by the end of the century whilst the bc ensemble projects an increase of 165 mm/yr. the bias-correction method applied here, clearly changes the climate signal from the combined gcm-rcm. this contrasts with other bias-correction methods, such as the delta change bias-correction, which has no such effect. pet projections follow a similar pattern as temperature, with larger increases projected by the bc models compared to the uncorrected projections, and with the largest increase by the end of the century. notably, the ensemble change for pet is always lower than the change projected for precipitation. table 1 shows the projected changes in mean annual temperature, pet and precipitation for individual models by the end of the century. clusters are observed, such as models that project a warmer (e.g. canesm2-remo2015 and all rcms driven by hadgem2-es under rcp 8.5) or a wetter climate (canesm2-remo2015, ipsl-cm5a-mr-rca4, hadgem2-es-hirham5 under rcp 8.5) compared to the ensemble mean. further clusters emerge among models that project an increase in water stress (where the increase in pet is larger than the increase in precipitation), such as hadgem2-es-cclm and hadgem2-es-remo2015 when driven by rcp 8.5. these clusters can provide insights into the impacts of climate change on danish water resources. uncertainty of the projections the ensemble spread from the bc simulations is smaller than the spread of the uncorrected models for temperature and pet when driven by rcp 8.5. for precipitation, the ensemble spread decreases for both rcps. the standard deviation of the mean annual precipitation from 2071 to 2100 is reduced by bias-correction from 166 mm to 122 mm for rcp 4.5 and from 211 mm to 139 mm for rcp 8.5. the spread or ‘uncertainty’ in projections comes from the choice of gcm, rcm or rcp and the natural variability expressed in the models. to assess the contribution of each source of uncertainty to the overall spread of projections, we analysed the signal-to-noise ratio (snr) of the precipitation and temperature projections driven by rcp 8.5 for the middle and end of the century (table 2). the snr of an ensemble is defined as the projected mean divided by the standard deviation of the ensemble. thus, a low snr implies that the uncertainty of the projection is high. our analysis has some limitations, which we acknowledge here. first, the full range of all possible combinations of gcms and rcms were not available for the uncertainty analysis. second, some of the available gcm-rcm combinations were run with different initial conditions and third, not all rcms are driven by the same gcms. considering these limitations, we used the gcm-rcm combinations driven by hadgem2-es to assess rcm uncertainty. gcm uncertainty was estimated by averaging the output of the remo2015 and rca4 rcms (each one driven by three different gcms). rcp uncertainty was evaluated using the gcm-rcm combinations available for both scenarios. uncertainty associated with natural variability was assessed using simulations with two different initial conditions (table 1). for temperature, the largest source of uncertainty in the uncorrected models is the choice of rcp scenario used. the uncertainty associated with natural variability is largest by the middle of the century and then reduces. finally, the uncertainty associated with the gcm is larger than that of the rcm, which represents the smallest source of uncertainty, overall. these results are similar to the findings of hawkins & sutton (2011) for projections of global mean temperature. for precipitation, the choice of gcm and rcp provides the largest sources of uncertainty by the middle of the century and the end of the century, respectively. the next largest source of uncertainty is the rcm followed by natural variability. hawkins & sutton (2011) estimated that the model uncertainty is larger than the uncertainty associated with table 2. signal to noise ratio for temperature (t) and precipitation (p) uncertainty source 2041–2070 2071–2100 2041–2070 2071–2100 2041–2070 2071–2100 2041–2070 2071–2100 gcm 5.7 5.7 1.2 2.0 6.3 6.3 2.0 2.7 rcm 15.3 19.9 2.4 2.8 9.9 10.1 1.8 2.5 rcp 2.6 2.6 1.7 1.4 3.2 3.0 2.1 1.6 nv 5.0 42.0 12.9 4.5 6.3 16.1 3.7 11.4 raw bc t (°c) p (mm) t (°c) p (mm) gcm: general circulation model. rcm: regional climate model. rcp: representative concentration pathway. nv: natural variability. e2019430102-05 the emission scenario, with little influence from natural variability. this agrees with our results, but in denmark, rcp becomes the largest source of uncertainty by the end of the century. bias-correction does not alter the uncertainty associated with the temperature projections. however, bias correction of the precipitation data causes the choice of rcm to become the largest source of uncertainty by the middle of the century, and the second largest source of uncertainty by 2100. spatial distribution of the projections precipitation is projected to increase throughout denmark, but the relative magnitude of this change varies according to location. the projected change in the uncorrected models ranges from +10% to +22% by the end of the century, compared to the 1981–2010 reference period (fig. 2a), whereas the bc projections range from +12% to +31% (fig. 2b). similarly, the standard deviation of the uncorrected projections varies between +3% and +19% and between +4% and +21% for the bc models. bias correction generally leads to even higher projections of precipitation by the end of the century. the standard deviation is less effected. the spatial distribution of change is relatively homogeneous over inland denmark. variations in the projections are mostly observed on the coast cells in both the uncorrected and bc models. however, after bias-correction this variation along the coast increases as indicated by the large standard deviation. this could be due to the interpolation method in the observation dataset, which lacks point data in the coast cells. outlook this study provides an overview of the bias-corrected projections from current state-of-the-art climate models, which were not previously available for denmark. by identifying the contribution of each uncertainty source and providing fig. 2. rcp 8.5 annual precipitation change (%) by the end of the century (2071–2100) relative to the 1981–2010 reference period for the a: uncorrected and b: bias-corrected ensemble. standard deviation for the c: uncorrected and d: bias-corrected ensemble. change (%) 0 < 10% 10-12% 12-14% 14-16% 16-18% 18-20% 20-22% > 22% st. dev. < 2% 2% -4% 4% -6% 6% -8% 8% -10% 10% -12% 12% -14% 14% -16% 16-18% > 18% a. uncorrected ensemble, mean change b. bias-corrected ensemble, mean change c. uncorrected ensemble, standard deviation d) d. bias-corrected ensemble, standard deviation e2019430102-06 *corresponding author: ernesto pasten-zapata | e-mail: epz@geus.dk 1 geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350, copenhagen k, denmark. the projected change from the ensemble and from each individual model, we provide a basis upon which to plan future assessments of the impacts of climate change on danish water resources. the data represent a useful input to the danish national water resources model (dk-model) for the analysis of climate change impacts. however, this initial analysis is aggregated for the whole of denmark and projections vary across the country. further research will focus on assessing monthly and seasonal changes in the projections as well as using these post-processed models to evaluate the projected impacts on danish hydrology. acknowledgements this research was funded by the aquaclew project and the danish agency for data supply and efficiency (the danish ministry of energy, utilities and climate). we thank dmi for providing the gridded observations and the euro-cordex models for the danish domain. references chen, j., brissette, f.p. & lucas-picher, p. 2015: assessing the limits of bias-correcting climate model outputs for climate change impact studies. journal of geophysical research: atmospheres 120, 1123–1136. https://doi.org/10.1002/2014jd022635 hawkins, e. & sutton, r. 2011: the 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https://doi.org/10.1002/2014jd022635 https://doi.org/10.1007/s00382-010-0810-6 https://doi.org/10.1007/s10113-013-0499-2 https://doi.org/10.1007/s10113-013-0499-2 https://doi.org/10.1002/joc.3518 https://doi.org/10.1175/jcli-d-12-00821.1 https://doi.org/10.1007/s40641-016-0050-x https://doi.org/10.1007/s40641-016-0050-x https://doi.org/10.1002/2014ef000259 https://doi.org/10.1017/9781107588783 https://doi.org/10.1038/nature08823 https://doi.org/10.1038/nature08823 https://doi.org/10.1016/j.jhydrol.2014.08.026 https://doi.org/10.5281/zenodo.2558193 https://doi.org/10.5281/zenodo.2558193 https://doi.org/10.1016/j.jhydrol.2013.02.015 https://doi.org/10.1016/j.jhydrol.2013.02.015 https://doi.org/10.1175/bams-d-11-00094.1 https://doi.org/10.1007/s10584-011-0148-z https://doi.org/10.1007/s10584-011-0148-z https://doi.org/10.34194/geusb-201943-01-02 e2019430203-01 titanite (catisio5) occurs as a rare mineral in magmatic and metamorphic rocks. it is commonly found in clastic sedimentary rocks as an accessory heavy mineral – a mineral of high density. recently, u-pb dating of single-grains of detrital titanite has been shown to be a useful tool in sedimentary provenance studies (e.g. mcateer et al. 2010; thomsen et al. 2015). titanite u-pb geochronologies can add important information to constrain the sediment sources of rocks and basins, and can help date precipitation of titanite. however, there are a number of complicating factors that must be taken into consideration for reliable application of titanite u-pb dating in provenance studies. first, titanite is less stable than zircon – the most commonly employed dating target. for example, in palaeocene sediments in the north sea, titanite rarely occurs as detrital grains at burial depths greater than 1400 m (morton 1984). it can also show dissolution features due to weathering and burial diagenesis (e.g. morton 1984; turner & morton 2007). second, titanite may precipitate during burial diagenesis, which would reflect the burial history of sediments and not their provenance. precipitation of authigenic titanite is documented from deeply buried (i.e. at temperatures greater than 100°c) volcaniclastic sandstones and mudstones (helmond & van de kamp 1984; milliken 1992) and intrusionassociated mineralisation in volcanic permian sandstones (van panhuys-sigler & trewin 1990). moreover, titanite also occurs in shallow-buried jurassic sandstones with no volcanic affinity (morad 1988). thus, the formation of titanite is not necessarily linked to a volcaniclastic source, but nevertheless, the presence of volcanic material seems to promote titanite precipitation. if authigenic titanite precipitation was incorrectly identified as detrital, this would have considerable implications for provenance investigations, as apparently titanite-rich source rocks would be wrongly inferred to be present in the sediment source area. here, we present examples from the kangerlussuaq basin in southern east greenland of what appeared to be detrital titanite. however, new u-pb dating reveals that the titanite formed authigenically, and hence contributes to the burial history, and not the provenance, of the sediments. geological setting the kangerlussuaq basin in southern east greenland was formed by the north atlantic opening during the midcretaceous, and filled by cretaceous and palaeogene mudu-pb dating identifies titanite precipitation in paleogene sandstones from a volcanic terrane, east greenland rikke weibel*1 and tonny b thomsen1 geus bulletin is free to individuals and institutions in electronic form. the author(s) retain copyright over the article contents. research article | open access geus bulletin vol 43 | e2019430203 | published online: 08 july 2019 https://doi.org/10.34194/geusb-201943-02-03 32°40° 69° 69° 32° kangerlussuaq blossev ille kyst greenland ice sheet den mar k st ra it sødalen tertiary basalt precambrian/caledonian basement palaeozoic to cretaceous sediments 100 km fig. 1. distribution of the palaeozoic to cretaceous sediments in the kangerlussuaq area, southern east greenland, and their close relation to the tertiary plateau basalt (modified after henriksen et al. 2008). https://doi.org/10.34194/geusb-201943-02-03 e2019430203-02 stone-dominated sediments (fig. 1; larsen et al. 1999). the type of deposition changed during the late cretaceous to sand-prone fluvial deposits (schjelderup member), which cover large areas and probably formed due to crustal doming prior to the onset of volcanism (larsen et al. 1999). several kilometres of late palaeocene to early eocene continental flood-basalts cover the cretaceous and palaeocene sedimentary succession (e.g. nielsen et al. 1981; peate et al. 2003). during breaks in the volcanic activity, siliciclastic deposition resumed in a combination of shallow marine and deltaic environments (peate et al. 2003; larsen et al. 2016). the kangerlussuaq basin contains six intrabasaltic sandstone units (fig. 2; larsen et al. 2016). the earliest intra-basaltic sandstone has the highest content of siliciclastic material, whereas the volcaniclastic contribution increases in the second, and is the main constituent in the succeeding four intrabasaltic sandstone units (fig. 2). methods the suband intra-basaltic sandstones were investigated in thin section by optical microscopy, supplemented by scanning electron microscopy (sem) of thin sections and rock chips to establish the petrographical and diagenetic relationship between titanite and other mineral phases. a phillips xl 40 sem was operated using the secondary electron (se), back-scattered electron (bse) and the energy dispersive x-ray spectroscopy (eds), which combined a thermo nanotrace 30 mm2 detector surface window and a pioneer voyager 2.7 10 mm2 window si(li) detector system. the electron beam was generated by a tungsten filament operating at 17 kv and 50–60 µa. computer-controlled sem (cc-sem) was used for heavy-mineral analysis. heavy minerals were separated into grains by crushing and ultrasound treatment. the grains were sieved, and the heavy minerals were concentrated from the 45–750 µm fraction by heavy-liquid separation using bromoform. the resulting concentrate was embedded in epoxy and polished. carbon-coated polished blocks were analysed by sem under similar conditions as the petrographical investigations. the number of measured grains was typically 1200. the data were recalculated using a method described by keulen et al. (2009). identification of minerals was based on semi-quantitative eds (energy dispersive x-ray spectroscopy) analysis, whereby minerals of similar chemical composition were grouped. the mafic silicates include chlorite, amphibole, pyroxene, tourmaline and olivine. the ilmenite group covers altered ilmenite grains with a tio2 content of up to 64%, whereas altered grains with a tio2 content of 64–90% is considered to be leucoxene. fragments containing titanite were recorded as titanite only if the composition was tio2 > 15%, cao > 12% and sio2 > 15%. for majorand trace-element analysis, glass discs were produced by fusing ignited powdered samples with lithium tetraborate in pt/au crucibles. most major elements were obtained from the glass discs by x-ray fluorescence (xrf), using a bruker s8 tiger wavelength dispersive multichannel xrf spectrometer equipped with a rh-anode x-ray tube. the elements na and cu were acquired by atomic absorption spectrometry (aas). trace elements and rare-earth elements were measured by solution-mode inductively coupled plasma mass spectrometry (icp-ms) using a perkin elmer sciex (elan 9000) icp-ms and small pieces of glass disc dissolved in a mixture of hcl and hno3. titanite u-pb dating was successfully carried out on one of two samples. the mineral grains were embedded in epoxy mounts and analysed using a nwr213 laser ablation system, coupled to an element2 sf-icp-ms. titanite grains were hand-picked under a binocular microscope from a heavy mineral concentrate obtained by a holman–wilfley water-shaking table. data were acquired by single-spot analysis bracketed by the gj-1 zircon standard (jackson et al. 2004). results were validated by analyses of natural titanite standards a1772 and a968 (provided by y. lahaye), and the plešovice zircon standard (slama et al. 2008) throughout the analysis sequence, all yielding age accuracies of < 3% (2σ) deviation from reference values. data processing was performed offline using the software iolite v. 2.5 (paton et al. 2010, 2011) with the vizualage data reduction scheme (petrus & kamber 2012). data were corrected for background, session drift and down-hole isotopic fractionation. a common pb correction usually needs to be applied for titanites. however, titanites with a high proportion of common pb, as indicated by low 206pb/204pb (average c. 18 ± 1) and 207pb/204pb ratios (average c. 14 ± 0.5), present a different situation. ludwig (1998) reports that in such samples, if the common pb ratio is invariant, any error in the isotope ratios assigned to the common pb will result in a consistent bias, rather than a random variation, of the calculated 206pb/238u and 207pb/206pb radiogenic ratios. thus, in a ‘‘semitotal– pb/u isochron’’ approach (tera and wasserburg 1972), the backgroundand session-drift-corrected ratios can be plotted on the tera-wasserburg concordia diagram without correction for common pb. if (and only if) the true 206pb/238u and 207pb/206pb radiogenic isotope ratios yield comparable, concordant ages, will the non-common-pb-corrected data be dispersed along a line whose intercept with the concordia curve defines the age of the samples (ludwig 1998). this is the case for the titanite grains in this study. therefore, the lower intercept age reported here for the titanites is not corrected for common pb – assuming that the lower intercept age represents the titanite age due to a specific geological event. results the suband intra-basaltic sandstones show variations in their detrital grain compositions and exhibit distinctly different cementing phases. the sub-basaltic sandstones consist mainly of quartz, minor amounts of feldspar with suborditio2 content (%) 0 1 2 3 4 5 x heavy mineral (%) 0 10 20 30 40 50 60 70 80 90 100 ilmenite leucoxene rutile titanomagnetite chromite phosphates zircon garnet titanite epidote amphibole and pyroxene b siliciclastic vs volcaniclastic detritus (%) 0 20 40 60 80 100 f e x c siliciclastic detritus volcaniclastic detritus unit d fig. 2. bulk rock tio2 content and previously published heavy mineral assemblages for suband intra-basaltic sandstones (larsen et al. 2016). unit b: subbasaltic sandstone of the schjelderup member. units c, x, d, e, f: intra-basaltic sandstones. siliciclastic detritus dominates the sub-basaltic sandstones whereas the volcaniclastic content increase upwards. titanite comprises a relatively large proportion of the heavy minerals in the earliest and second intrabasaltic sandstones, even though the tio2 content is even higher in the later intra-basaltic sandstones. e2019430203-03 nate mica, rock fragments and heavy minerals. besides abundant quartz, the intra-basaltic sandstones are characterised by a high content of detrital feldspar and volcanic rock fragments. the authigenic phases in the sub-basaltic sandstones are dominated by quartz overgrowths and illite. the intrabasaltic sandstones are instead characterised by abundant authigenic chlorite, calcite, common feldspar and rare laumontite cement. the ti-rich heavy minerals in the sub-basaltic sandstones are mainly ilmenite, titanomagnetite, leucoxene, rutile and rare titanite (unit b; fig. 2). in contrast, titanite is the dominant ti-rich mineral in the intra-basaltic sandstones. anatase is the most abundant authigenic phase in the subbasaltic sandstones, whereas authigenic titanite is dominant in the intra-basaltic sandstones. titanite commonly occurs as single grains and is a common constituent of rock fragments. titanite forms tiny crystals together with other authigenic phases such as chlorite and/or calcite, possibly replacing volcanic glass fragments. authigenic titanite occurs as replacement of detrital ilmenite along fractures and ilmenite crystal rims and as authigenic precipitates, similar to chlorite, between the cleavage planes in mica (fig. 3). the bulk rock tio2 content increases with abundance of volcaniclastic material, and hence also upwards in the succession of intra-basaltic sandstones (fig. 2). although titanite is generally common in the heavy mineral assemblages, it only makes up a relatively small proportion in the volcaniclastic dominated intra-basaltic sandstones, due to more abundant mafic minerals. the titanite age data are plotted on a tera-wasserburg concordia diagram (fig. 4). the unanchored lower intercept age is reasonably well constrained due to the large spread in radiogenic pb/common pb ratios of the titanite grains. in one sample, an analysis of 156 titanite grains yielded a u-pb lower intercept age of 49 ± 2 ma (2σ) with a mean square of weighted deviates (mswd) of 9.6 on both the tera-wasserburg and conventional (wetherill) concordia diagrams. discussion the ti-rich minerals likely derived from the regional gneiss basement that is assumed to have formed contemporarily with the crystalline basement in the scoresby sund region, north of the kangerlussuaq area, and which yielded u-pb zircon ages of 2600 to 3000 ma (henriksen et al. 2008). this is supported by a u-pb zircon age of 2700–3700 ma from a paleogene sandstone in the kangerlussuaq area (whitham et al. 2004). our data show that replacement of ilmenite and titanomagnetite by titanite, and titanite precipitation occurred 5 µm 10 µm 5 µm b a c ti ti ilm ti ch ch ti pl ab fig. 3. micrographs of authigenic titanite. a: titanite (ti) partially replacing an ilmenite (ilm) grain (intra-basaltic sandstone; bse micrograph). b: titanite and chlorite (ch) precipitated between the cleavage planes in mica (intra-basaltic sandstone; se micrograph). c: titanite intergrown with albite (ab) and ca-rich plagioclase (pl; intra-basaltic sandstone; se micrograph). e2019430203-04 49–42 ma (fig. 4). this probably coincides with maximum burial of the sediments, as thick (6 to 8 km) flood basalt units were extruded over the kangerlussuaq area with the main eruption phase at 60–50 ma (nielsen & brooks 1981; larsen & tegner 2006; brooks 2011). furthermore, maximum burial must have occurred prior to the first cooling episode during the late eocene (40–35 ma) as recorded by apatite fission track analyses (japsen et al. 2014). the reason for preferential precipitation of anatase in subbasaltic sandstones and titanite in intra-basaltic sandstones, could be that titanite is favoured either by (1) higher temperatures or (2) liberation of ca and si simultaneously with alteration of fe-ti oxides in intra-basaltic sandstones. the abundance of titanite increases with the volcaniclastic input, hence the fifth and sixth intra-basaltic sandstones (units f, g; fig. 2) in the kangerlussuaq area, which are almost completely dominated by volcaniclastic material, show more abundant titanite than the lower intra-basaltic sandstones, which are comprised of mixed siliciclastic and volcaniclastic material (units c, x; fig. 2). previous investigations show that titanite precipitation is associated with the presence of volcanic rock fragments (helmond & van de kamp 1984; van panhuys-sigler & trewin 1990) possibly because ca, si and ti are likely liberated concurrently during alteration of volcanic rock fragments. formation water from adjacent volcanic rocks could similarly have contributed elements for titanite precipitation. here, the effect of temperature can be disregarded since the suband intrabasaltic sandstones have both experienced similar burial histories of up to 6–8 km burial depth. alternatively, titanite might have precipitated as a result of the higher heat flux from intrusions or extruded lava, but in this case, a higher abundance of titanite would be expected immediately adjacent to the lava piles or intrusions, which we did not observe. conclusions the suband intra-basaltic paleogene sandstones from the kangerlussuaq area show a major difference in the dominant ti-bearing phases. detrital rutile, ilmenite and leucoxenereplaced fe-ti oxides dominate in the sub-basaltic sandstones and anatase is a common authigenic phase. in the intra-basaltic sandstones, titanite is the dominant ti-phase and here it replaces detrital ti-rich grains and precipitates as tiny crystals. this reflects different diagenetic changes and not a shift in provenance. u-pb dating of titanite documents that titanite formed during diagenesis, c. 49 ma, at maximum burial. despite similar burial history, different diagenetic paths are probably caused by the absence or presence of volcanic material in the suband intra-basaltic sandstones, respectively. seemingly detrital titanite is in fact titanitereplaced fe-ti oxides, and hence does not originate from the sediment source area. care must be taken when working with sediment where abundant volcanic material is present, since it is interpreted to have caused the major difference in dominant ti-phase between the suband intra-basaltic sandstones. the presence of partly titanite-replaced detrital fe-ti oxides may indicate that all titanite is authigenic and u-pb dating may be necessary to establish its true origin. acknowledgments reviewer andrew morton and david chew are thanked for constructive comments, which improved the paper. 1.2 lower intercept = 49.7 ± 0.6 ma 1.0 0.8 0.6 0.4 0.2 0.0 0 50 100 ma 238u / 206pb 20 7 p b / 20 6 p b 200 ma 4900 ma 4800 ma 4700 ma 4600 ma 4500 ma 4100 ma 3500 ma 3000 ma 2500 ma 2000 ma 1000 ma 500 ma 50 ma 100 150 fig. 4. tera-wasserburg diagram. the lower intercept age of c. 49 ± 2 ma calculated for sample 514625 (comprising 158 la-icpms analyses) is indicated. titanites were not corrected for common pb content. this age most likely represents titanite precipitation and simultaneous replacement of fe-ti oxides during maximum burial due to overlying 56–60 ma tertiary flood-basalts. e2019430203-05 *corresponding author: rikke weibel | e-mail: rwh@geus.dk 1 geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350, copenhagen k, denmark. references brooks, c. k. 2011: the east greenland rifted volcanic margin. 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(eds): heavy minerals in use. developments in sedimentology, elsevier 58, 393–412. https://doi.org/10.1016/s0070-4571(07)58014-3 van panhuys-sigler, m. & trewin, n.h. 1990: authigenic sphene cement in permian sandstones from arran. scottish journal of geology 26, 39–144. https://doi.org/10.1144/sjg26020139 whitham, a.g., morton, a.c. & fanning, c.m. 2004: insights into cretaceous–palaeogene sediment transport paths and basin evolution in the north atlantic from a heavy mineral study of sandstones from southern east greenland. petroleum geoscience 10, 61–72. https://doi. org/10.1144/1354-079302-506 how to cite weibel, r. & thomsen, t.b. 2019: u-pb dating identifies titanite precipitation in paleogene sandstones from a volcanic terrane, east greenland. geological survey of denmark and greenland bulletin 43, e2019430203. https://doi.org/10.34194/geusb-201943-02-03 e2019430203-06 mailto:rwh%40geus.dk?subject= https://doi.org/10.1306/m37435c15 https://doi.org/10.1016/j.gloplacha.2014.01.012 https://doi.org/10.1016/j.gloplacha.2014.01.012 https://doi.org/10.1016/j.gexplo.2009.04.001 https://doi.org/10.1016/j.gexplo.2009.04.001 https://doi.org/10.1016/j.lithos.2006.03.032 https://doi.org/10.1016/j.lithos.2006.03.032 https://doi.org/10.1144/0050337 https://doi.org/10.1144/0050337 https://doi.org/10.1144/pgc8.13 https://doi.org/10.1016/s0016-7037(98)00059-3 https://doi.org/10.1016/s0016-7037(98)00059-3 https://doi.org/10.1016/j.precamres.2010.05.013 https://doi.org/10.1306/d42679dd-2b26-11d7-8648000102c1865d https://doi.org/10.1306/d42679dd-2b26-11d7-8648000102c1865d https://doi.org/10.1016/0037-0738(88)90016-4 https://doi.org/10.1016/0037-0738(88)90016-4 https://doi.org/10.1180/claymin.1984.019.3.04 https://doi.org/10.1180/claymin.1984.019.3.04 https://doi.org/10.1144/gsjgs.138.5.0559 https://doi.org/10.1144/gsjgs.138.5.0559 https://doi.org/10.1029/2009gc002618 https://doi.org/10.1039/c1ja10172b https://doi.org/10.1144/0016-764902-071 https://doi.org/10.1111/j.1751-908x.2012.00158.x https://doi.org/10.1111/j.1751-908x.2012.00158.x https://doi.org/10.1016/j.chemgeo.2007.11.005 https://doi.org/10.1016/0012-821x(72)90128-8 https://doi.org/10.1016/0012-821x(72)90128-8 https://doi.org/10.1016/s0070-4571(07)58014-3 https://doi.org/10.1144/sjg26020139 https://doi.org/10.1144/1354-079302-506 https://doi.org/10.1144/1354-079302-506 https://doi.org/10.34194/geusb-201943-02-03 geological survey of denmark and greenland bulletin 41, 2018, 21-24 21 a glacitectonite is defined as a brecciated sediment or a cataclastic sedimentary rock formed by glaciotectonic deformation (pedersen 1988). the term tectonite was initially introduced by sander (1912), mainly for tectonically brecciated metamorphic rocks in the alps. in the classic work on cataclastic rocks, higgins (1971) stated that the term covered all rocks with fabric displaying coordinated geometric features related to continuous flow during deformation.therefore brecciated lithologies formed by glaciotectonic deformations can be termed tectonites. banham (1977) suggested the prefix glacito clarify the relation to glacial dynamics. furthermore, pedersen (1988) suggested the application of the bedrock prefix. thus, a chalk-glacitectonite is a brecciated chalk formed by shear deformation during a glacial advance over an exposed bedrock surface of chalk (fig. 1). hence the term describes a sedimentary rock in which the primary structures are so disturbed that they cannot be continuously traced, and a glacitectonic fabric developed as joint fractures or shear surfaces superimposed on the lithology. the significance of recognising chalk-glacitectonite from chalk and limestone bedrock is the difference in textural properties, which is fundamental in geological modelling. in areas dominated by glaciotectonic complexes, which include thrust sheets of pre-glacial sedimentary rocks, the sheets are subject to shearing and dragged along the sole of the ice during its movement over the glaciotectonic complex. due to truncation and shear-drag, the glacitectonite forms at the base of the deformational layer in a lodgement till. from the source area, which typically is a detachment anticline, the chalk-glacitectonite, an important lithology in former glaciated terrains covering chalk and limestone bedrock stig a. schack pedersen, peter gravesen and klaus hinsby 5 4 3 2 1 in cr ea sin g gla cio te ct on ic sh ea r d ef or m at io n lo dg em en t til l chalk or limestone limy till fig. 1. five steps in the progressive formation of chalk glacitectonite and limy till developed from bedrock of danian limestone. the example illustrates the variation of deposits differentiated in the geological mapping of north-earstern djursland, central denmark (from pedersen & petersen 1997). 1: undisturbed danian limestone occurring in the lower part of the coastal cliff at sangstrup klint. 2: anastomosing jointing is found in the limestone in the upper part of the cliff exposure. note that the smallest angle between joints is located with a half-angle divide in the horizontal plan. this corresponds to a lateral stress in the foreland to an advancing ice margin. 3: clasts of chalk have been broken off and displaced in a fine-grained matrix; a chalk-glacitectonite is formed. 4: during increased shearing the chalk clasts become more and more crushed with chalk pieces floating in a chalk-clay matrix. in-basinal erratics comprise clasts of danian limestone and flint, ex-basinal erratics include basement stones (gneiss and granite), which start to appear in the glacitectonite derived from the overlying lodgement till. 5: during the continuous translocation away from the source area the chalk-glacitectonite is transformed into limy till (chalk moraine), which may also be classified as a local till dominated by in-basinal clasts of chalk and flint. fig. 2. a one-metre thick chalk-glacitectonite exposed in a cliff section in the northern part of stevns klint displays shear banding of clayey till material with cataclasts of chalk and flint. the source area for the chalk is danian limestone which occurs more than 500 m from the exposure. © 2018 geus. geological survey of denmark and greenland bulletin 41, 21–24. open access: www.geus.dk/bulletin http://www.geus.dk/bulletin 2222 glacitectonite thins out in the direction of transport from 1–2 m (fig. 2) to a thin shear zone only a few centimetres thick over a distance of one to a few kilometres (pedersen 1996). moreover, brecciation of thrust sheets displaced by glacial thrusting occurs within glaciotectonic complexes. the deformation ranges from initially anastomosing jointing (figs 1, 3) to brecciation with bedrock clasts in crushed bedrock matrix (fig. 4). the tectonic breccia distributed from the décollement zone at the base to the truncating glacial unconformity at the top may additionally be termed glacitectonites. here we describe the occurrence and identification of chalk-tectonites. occurrences of chalk-glacitectonites the occurrences of chalk-glacitectonites are naturally related to the areas dominated by bedrock of chalk and limestone which in denmark includes the eastern, north-eastern and northern regions (fig. 5). bedrock exposures are found at møns klint and stevns klint in eastern denmark, sangstrup klint (ne djursland) in central denmark, in the chalk pits in aalborg and in limestone pits in adjacent areas in ne himmerland. chalk-glacitectonites occur at these outcrops. furthermore, the cliffs at agger, bulbjerg and hanstholm in nw jylland show outcrops of chalk and limestone. in addition, chalk that appears in the aureole of salt structures at, for example, gassum, suldrup, batum, erslev, uglev and thisted represent potential areas of glacitectonite occurrences. the relation between the overburden of quaternary deposits and the formation of glacitectonites is independent of the depth of the deposits. thus a glacitectonite should always be expected between the top of the chalk and an overlying till. however, the till and glacitectonite may have been removed by erosion. identification of chalk-glacitectonites there is a general understanding of the complexity of hydraulic properties related to areas with limestone and chalk located at shallow depths below quaternary overburden (downing et al. 1993). this is e.g. recognised in the greater copenhagen area where groundwater flow paths are difficult to predict and the permeability in the glacially disturbed chalk layers of the københavn kalk formation and the underlying danian bryozoan limestone are notably higher than in the underlying undisturbed limestones (klitten et al. 2006; bonnesen et al. 2009; galsgaard et al. 2014). research into the difficulty in predicting groundwater flow paths in shallow chalk aquifers is conduceted in an on-going eu project investigating fig. 3. the initial glaciotectonic deformation is a low-angle, anastomosing jointing, which is illustrated by an example of fractured cretaceous chalk exposed in the northern part of the stevns klint cliff section. fig. 4. a chalk-glacitectonite developed with a limy matrix and rotated chalk clasts. thin dark clayey shear bands illustrate the substantial amount of displacement within the rock type. detail from the cliff section at hvide klint, south coast of møn. eocene–miocene paleocene above danian danian limestone upper cretaceous chalk older than upper cretaceous 50 km mk stk sak su d a hi aa jylland t h b e g fa u fig. 5. geological map showing the distribution of chalk and limestone in the bedrock of denmark. modified from håkansson & pedersen (1992). h: hanstholm. b: bulbjerg. t: thisted. a: agger. e: erslev. u: uglev. aa: aalborg. su: suldrup. hi: himmerland. g: gassum. d: djursland. sak: sangstrup klint. stk: stevns klint. mk: møns klint. fa: falster. 23 subsurface water technologies to control saltwater intrusion (zuurbier et al. 2016) on southern falster, se denmark. the project focuses on the impact of climate change on the salinity of the groundwater resources (rasmussen et al. 2013). at the study site the top surface of the upper cretaceous chalk is situated at about 10 m below the surface; it is overlain by a 5 m thick unit of glacial sediments and 5 m marine sand (fig. 6). however, at a depth of 15 to 18 m there is a layer of chalk with gravel and pebbles of basement rocks. based on an evaluation of data from other wells in the area, it became evident that another zone with basement gravel and pebbles existed even deeper at a level from 30 to 40 m below the surface. these findings have implications for the understanding of groundwater flow around the wells as the complexity of the hydraulic characteristics markedly changes the aquifer’s behaviour. a model of the glacitectonite occurrence was established based on a glaciodynamic concept of the area (fig. 6). a resistivity log from a nearby well supported the model predictions with a glacitectonite on top of the undisturbed chalk (pedersen & hinsby 2017). on-going studies indicate that in some parts of the chalk reservoir the transmissivity behaves as single porosity aquifers, while other parts behave like fractured dual porosity aquifers. g lac io te ct on ic de pr ess ion m ar gin al m or ain e 2 km g lac io te ct on ic de pr ess ion m ar gin al m or ain e bøtø ringvej well location a 242.394 242.384 18 31 35 43 10 d ep th b el ow su rfa ce (m ) c on ce ale d m el tw at er c ha nn el chalk glacitectonite décollement zone chalk bedrock w e breciated chalk 20 m overburden of postglacial marine sand cover of glacial deposits on chalk top surface c fig. 6. the hydrogeological investigation site on falster (the bøtø case): a: geological map of the area demonstrating the glacial geological setting. b: two borehole logs demonstrating the lithological settings. c: block diagram illustrating the features and glacitectonites in the upper cretaceous beds. fig. 7. an about one-metre thick bed of chalk-glacitectonite separates two till beds exposed at the north coast of stevns, se denmark. 0 5 10 15 20 25 30 35 d ep th b el ow su rfa ce (m ) 38 dgu well no 242.394 dgu well 242.395 0 5 10 15 20 soil marine sand peat and gyttja glaciofluval sand clayey lodgement till cretaceous chalk flint in chalk chalk-glacitectonite 21 d ep th b el ow su rfa ce (m ) b 2424 the position of glacitectonites in the glaciodynamic development of the quaternary successions the chalk-glacitectonites occur basically at two different positions in the glaciodynamic sequence: either as tectonic breccias on top of chalk bedrock, or as shear translocated chalk debris at the sole of a basal till. in the first position the chalk-glacitectonite may be difficult to distinguish from undeformed bedrock. this is especially the case with identification of lithologies from drill-hole samples. identification requires that small impurities, basement pebbles etc., displaced into the fractures, are recognised and documented. the second position of chalk-glacitectonites is easy to recognise due to the unmistakable variation in lithology (pedersen & gravesen 2016; fig. 7). the bedrock material appears in a succession of glacial deposits. the typical glaciodynamic sequence contains a meltwater unit of clay/silt grading up into sand coarsening up into glaciofluvial gravel, eventually with a stone-bed of ice-contact deposits mirroring the proglacial environment. on top of the glaciofluvial sediments the basal till demonstrates the ice advance over the foreland. the till is divided into a basal deformational layer and an upper lodgement layer. thus the chalk-glacitectonite, representing the deformational layer, documents the transition from the foreland setting to the subglacial setting. final remarks chalk-glacitectonites are an important lithology to be identified in glacial terrains with bedrock comprising chalk and limestone, i.e. where the pre-quaternary surface consists of limestones and related carbonate rocks. chalk-glacitectonites are divided into two main types based on the structural setting in a glaciotectonic complex: (1) brecciated sedimentary rocks deformed within the stratigraphic succession of the deformed bedrock, and (2) brecciated rock deformed below a basal till and shear-mixed into the lodgment till. the recognition of chalk-glacitectonites is important for geological and groundwater-flow modelling addressing hydrogeological and geotechnical problems. due to the glacial deformation these sedimentary rocks are expected to show higher permeability than undeformed bedrock. acknowledgement this study was part-funded by the eu horizon2020-project ‘subsol’ (grant agreement no. 642228, www.subsol.org). references banham, p.h. 1977: glacitectonites in till stratigraphy. boreas 6, 101–105. bonnesen, e., larsen, f., sonnenborg, t.o., klitten, k. & stemmerik, l. 2009: deep saltwater in chalk of north-west europe: origin, interface characteristics and development over geological time. hydrogeology journal 17, 1643–1663.  downing, r.a., price, m. & jones, g.p. (eds) 1993: the hydrogeology of the chalk of north-west europe, 310 pp. oxford: oxford university press. galsgaard, j., rhode, r., jakobsen, r. & jakobsen p.r. 2014: strømning og stoftransport i kalklagene på den københavnske vestegn. geologisk og hydrogeologisk vidensopsamling og typemodel. geo projekt 37208, rapport 1, 87 pp. higgins, m.w. 1971: cataclastic rocks. u.s. geological survey professional paper 687, 97 pp. håkansson, e. & pedersen, s.a.s. 1992: geologisk kort over den danske undergrund, 1:500 000, map sheet. copenhagen: varv. klitten, k., larsen, f. & sonnenborg, t.o. 2006: saltvandsgrænsen i kalkmagasinerne i nordøstsjælland, hovedrapport, 45 pp. copenhagen: geological survey of denmark and greenland & institut for miljø og ressourcer. pedersen, s.a.s. 1988: glacitectonite: brecciated sediments and cataclastic sedimentary rocks formed subglacially. in: goldthwait, r.p. & matsch, c.l. (eds): genetic classification of glacigenic deposits, 89–91. rotterdam: a.a. balkema. pedersen, s.a.s. 1996: progressive glaciotectonic deformation in weichselian and palaeogene deposits at feggeklit, northern denmark. bulletin of the geological society of denmark 42, 153–174. pedersen, s.a.s. & gravesen, p. 2016: risikovurdering af skredforhold langs stevns klint. danmarks og grønlands geologiske undersøgelse rapport 2016/33, 87 pp. pedersen, s.a.s. & hinsby, k. 2017: foreløbige resultater af boringer ned i skrivekridt ved bøtø ringvej, væggerløse, falster. danmarks og grønlands geologiske undersøgelse rapport 2017/28, 14 pp. pedersen, s.a.s. & petersen, k.s. 1997: djurslands geologi, 96 pp. copenhagen: danmarks og grønlands geologiske undersøgelse. rasmussen, p., sonnenborg, t.o., goncear, g. & hinsby, k. 2013: assessing impacts of climate change, sea level rise, and drainage canals on saltwater intrusion to coastal aquifer. hydrology and earth system sciences 17, 421–443. sander, b. 1912: über einige gesteinsgruppen des tauernwestendes. jahrbuch der geologischen reichanstalt 62, 219–288. zuurbier, k. et al. 2016: how subsurface water technologies (swt) can provide robust, effective, and cost-efficient eolutions for freshwater management in coastal zones. water resources management 31, 671–687. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen, denmark. e-mail: sasp@geus.dk. http://www.subsol.org mailto:sasp@geus.dk geological survey of denmark and greenland bulletin 28, 2013, 45-48 45 the norite belt in the mesoarchaean maniitsoq structure, southern west greenland: conduit-type ni-cu mineralisation in impact-triggered, mantle-derived intrusions? adam a. garde, john pattison, thomas f. kokfelt, iain mcdonald and karsten secher with the recent discovery of the giant, deeply eroded, 3 ga maniitsoq impact structure in southern west greenland (garde 2010), an enigmatic, c. 75 by 15 km large, curvilinear belt of undeformed norite intrusions with ni-cu mineralisation was re-interpreted as representing crustally contaminated melts derived from the mantle in the wake of the impact (fig. 1; garde et al. 2012). the norite belt (nielsen 1976; secher 1983) was discovered in the early 1960s by the mining and exploration company kryolitselskabet øresund a/s, and more than one hundred shallow exploration holes were drilled by the company in the period 1965–1971. the mineralisation has subsequently been investigated by cominco ltd., falconbridge ltd. and nunaminerals a/s. in 2011, the re-interpretation of the norite belt, and recent availability of improved airborne geophysical exploration tools, prompted the canadian company north american nickel inc. (nan) to resume exploration. © 2013 geus. geological survey of denmark and greenland bulletin 28, 45–48. open access: www.geus.dk/publications/bull d = 100 km d = 50 km d = 150 km 2 5 4 d = 100 km ‘spotty hill’ ‘imiak hill’ sillisissannguit nunaat ‘fossilik’ 65°n 10 km maniitsoq ala nn gu a toqqusap nunaa ? ? ? 52°w greenland jurassic / archaean carbonatite surficial cover c. 3 ga cataclasis >3 ga rocks palaeoproterozoic mafic dykes norite belt, ni mineralisation postkinematic diorite intrusions granitic plutons archaean, c. 3 ga rocks approx. melt zone (extent revised from garde et al. 2012) metavolcanic belts and related rocks finnefjeld domain (cataclastic rocks) ttg orthogneiss fig. 1. geological map of the central part of the maniitsoq impact structure with the cataclastic finnefjeld domain in its centre, currently known distribution of norites and postkinematic diorites, and three labelled diamond drilling sites from 2012. reference circles with 50 and 100 km diameters are shown. 4646 this contribution gives a short description of the norite belt, outlines its ni-cu mineralisation and presents a preliminary interpretation. it is well known that the 1.85 ga sudbury impact structure in canada hosts a group of worldclass ni-cu deposits at the base of the impact melt sheet (eckstrand & hulbert 2007). the ni-cu mineralisation of the norite belt in west greenland is profoundly different as it is located in lower-crustal mafic intrusions, but we suggest here that also this was governed by impact-related magmatic processes. the maniitsoq structure the maniitsoq structure (garde et al. 2012) constitutes the deeply eroded, deep-crustal remains of a giant extraterrestrial impact that struck an active magmatic arc 3 ga ago in the north-central part of the north atlantic craton (fig. 1). the visible remains of the structure comprise a core measuring 35 by 50 km of thoroughly crushed, mechanically mixed and subsequently welded fragments of orthogneiss and amphibolite (together constituting the finnefjeld domain), surrounded by an up to c. 75 km wide, annular melt zone variably affected by both crushing and melting, and a peripheral, less damaged zone with a diameter of approximately 150 km. large parts of the entire structure have been affected by intense hydrothermal alteration and related, fluid-induced, partial melting. these features, and their large geographical extent, suggest that the original maniitsoq crater was larger than all previously known terrestrial impact structures. the norite belt and postkinematic diorites the norite belt (see above) and associated postkinematic diorites in the southern part of the impact structure (berthelsen 1960; garde 1991) collectively comprise numerous mafic to ultramafic intrusions ranging in size from small dykes to inclined, elongate bodies up to several square kilometres in outcrop size, hosted by mesoarchaean orthogneisses and mafic, metavolcanic belts. they were emplaced close to 3.0 ga ago (see garde et al. 2012), and their undeformed state and occasional xenoliths of crushed orthogneiss and fluidised microbreccia document that they post-date the impact. the intrusions are locally cut by few centimetres thick, white, undeformed pegmatites. most intrusions display inclined hostrock contacts, and the exposed bodies may thus represent open-ended magma conduits. the hybridised contact zones are diffuse, up to about 10 m wide and commonly hydrated (fig. 2). the known extent of the norite belt is shown on fig. 1, but at least a few additional, unmapped bodies are known to occur east and west of the main belt. the postkinematic diorites have a more random distribution and extend up to c. 70 km south and east of the centre of the impact structure. they have only been mapped systematically south of 65°n. the noritic and dioritic rocks typically form dark brown, crumbly weathering, mediumto coarse-grained, mostly structureless masses. they largely consist of variable proportions of orthopyroxene and andesine plagioclase with mediumto coarse-grained, granular textures. rhythmic magmatic layering is rare, but has been observed, e.g. at sillisissannguit nunaat (fig. 1; secher 1983). both small and large intrusions locally display proto-orbicular quench textures comprised of skeletal, rounded to ellipsoidal plagioclase crystals up to about 10 cm in size, and interstitial orthopyroxene (fig. 3), and locally display metre-scale layering. bulk compositions cover a large range from dioritic to fig. 2. typical homogeneous, undeformed norite outcrop with diffuse, hybridised contact zones to quartzo-feldspathic orthogneiss. from garde et al. (2012). fig. 3. proto-orbicular texture in norite, with rounded, skeletal plagioclase crystals and interstitial hypersthene. the hammer is 45 cm long. 47 noritic, but display a broadly uniform trace element distribution pattern; high mgo, cr and ni contents (≤21 wt%, 3685 ppm and 909 ppm, respectively) coupled with low nb and ta suggest an ultramafic precursor affected by abundant crustal contamination (garde 1991; garde et al. 2012). ni-cu mineralisation more than 50 sulphide-mineralised norite localities with uniform characteristics have been investigated over time (nielsen 1976; unpublished company reports). in typical, two-dimensional outcrops, the mineralised rocks form isolated, rusty weathered spots and lenses that rarely exceed 25 m in length (secher 1983). the sulphide paragenesis in fresh samples is almost invariably pyrrhotite-pentlanditechalcopyrite-pyrite, with accessory magnetite and ilmenite. the sulphide minerals form interconnected, semi-massive networks with inclusion-bearing, breccia-like textures characterised by rounded, centimetre-sized and larger lumps of the host rock (fig. 4). these textures closely resemble those found in magmatic, conduit-type ni deposits such as noril’sk in russia and lynn lake and voisey’s bay in canada (eckstrand & hulbert 2007). all significant nickel occurrences discovered to date at maniitsoq are entirely contained within norite intrusions. based on nan’s diamond drilling, combined with historical drill-hole information and three-dimensional modelling of airborne geophysical and down-hole electromagnetic data, the mineralisation can typically be described as steeply to vertically dipping pipeand sheet-like bodies comprised of heavily disseminated to near massive sulphide, often with abundant 0.2–5.0 cm subangular to subrounded inclusions of norite. in most cases, strongly mineralised zones are surrounded by a broad halo of weaker, disseminated mineralisation; however, in some places the contact between near solid sulphide and barren norite host rock is extremely sharp. diamond drilling in 2012 was performed at ‘imiak hill’, ‘spotty hill’ and ‘fossilik ii’ (cf. fossilik in fig. 1); the following, generalised figures of metal contents are solely intended to characterise the type of mineralisation without addressing economic aspects. the mineralisation has a high, uniform ni tenor averaging 6–9% ni recalculated to 100% sulphide, both at localities previously drilled by kryolitselskabet øresund a/s and also in new cores drilled by nan at ‘imiak hill’ and ‘spotty hill’ in 2012. typical ni contents in wt% of the total rock vary between 1 and 2% over several to many metres, with additional 0.1–0.6% cu, 0.01–0.07% co, up to 0.2 ppm pt and pd, and au in the ppb range (data from www.northamericannickel.com and secher 2001). these relatively low platinum-group element tenors are similar to the voisey’s bay deposit (lightfoot et al. 2012). discussion and interpretations origin of the noritic and dioritic intrusions garde (1991) and garde et al. (2012) interpreted the noritic and almost all of the dioritic intrusions as highly unusual, mantle-derived, crustally contaminated ultrabasic melts. this interpretation is based on (1) the wide, hybrid contact zones and (2) proto-orbicular interior textures of skeletal plagioclase and orthopyroxene, showing that the magmas were chilled by their hosts and crystallised rapidly (and thus indicating a very high emplacement temperature), (3) highly variable chemical compositions combined with (4) broadly similar overall element spectra consistent with mantle-crust mixtures, (5) ultramafic trace element signatures with very high cr, ni and v contents, and (6) mantle-like, strongly suprachondritic ru/ir, rh/ir, pt/ir and pd/ir ratios similar to high-mg basalts. fig. 4. inclusion-bearing, breccia-like texture of ni-cu mineralisation at ‘imiak hill’ (north american nickel inc. 2012 drill core mq-12-002 at 66.55 m). fig. 5. magmatic open-conduit model for the ni-cu mineralisation in the norite belt. schematic vertical section. modified from maier et al. (2001). ni-cu sulphides norite orthogneiss-amphibolite complex upper chamber lower chamber present-day erosion level 4848 interpretation of the ni-cu (-pge) mineralisation magmatic ni-cu-pge deposits are governed by several different factors (e.g. li et al. 2001; maier et al. 2001) including a fertile source (typically the mantle), where a significant ni contribution comes from interstitial sulphide minerals in addition to olivine. a high degree of melting is also required in order to ensure a complete dissolution of the sulphides with their metals into the primary magma. on the ascent of the magma, segregation of immiscible, sulphidic melt from the silicate melt must then occur in order to recapture the chalcophile elements of interest. in general terms, such segregation in response to sulphur saturation can be brought about by either fractional crystallisation in large magma chambers (potentially leading to pge-dominated deposits), or by crustal contamination in dynamic magma channels of restricted volume, potentially leading to ni-cu-dominated deposits (li et al. 2001). in the latter case, a suitable physical mechanism to separate and concentrate the exsolved, high-density sulphide liquid from the flowing magma is also required. the essential features of the sulphide occurrences in the norite belt are all compatible with an interpretation as magmatic, open-source mineralisation systems. these features include high-temperature ultramafic host rock melts, abundant field and chemical evidence of crustal contamination, inclusion-bearing, breccia-like textures in the mineralised rocks, a high ni tenor in the sulphides, and ni/cu and (ni + cu)/pge ratios characteristic of this type of mineralisation. the three-dimensional shapes of the mineralised volumes and mechanism(s) of physical extraction of the exsolved sulphide melt phase are currently poorly known, but a schematic section of a possible scenario is shown in fig. 5. concluding remarks widespread ultramafic magmatism, such as in the norite belt, does not belong in convergent orogenic systems (garde 1991; garde et al. 2012), although ultramafic intrusions can occur under special plate-tectonic circumstances of ridge subduction or delamination of the lower crust, for which there is no evidence at maniitsoq. furthermore, the protoorbicular textures and very high degrees of contamination in the norite belt and postkinematic diorites are highly unusual even for ultramafic melts. the possibility of impact-induced mantle melting associated with giant impacts has previously been discussed in the literature (jones et al. 2002), but no examples have been found prior to the case at maniitsoq (see garde et al. 2012). we conclude that the ni-cu occurrences in the norite belt are likely an impact-induced mineralisation. however, unlike sudbury they are not part of a differentiated impact melt sheet, but constitute a special variety of open-conduit systems, which are otherwise known from magmatic settings of non-impact origin. references berthelsen, a. 1960: structural studies in the pre-cambrian of western greenland. ii. geology of tovqussap nunâ. bulletin grønlands geologiske undersøgelse 25, 223 pp. eckstrand, o.r. & hulbert, l.j. 2007: magmatic nickel-copper-platinum group element deposits. in: goodfellow, w.d. (ed.): mineral deposits of canada: a synthesis of major deposit types, district metallogeny, the evolution of geological provinces, and exploration methods. geological association of canada special publication 5, 205–222. garde, a.a. 1991: post-kinematic diorite intrusions in archaean basement rocks around outer fiskefjord, southern west greenland. bulletin of the geological society of denmark 39, 167–177. garde, a.a. 2010: the 2975 ma maniitsoq impact structure in west greenland: the oldest and most deeply exposed meteorite crater on earth. abstracts and proceedings of the geological society of norway 1, 57–58. garde a.a., mcdonald, i., dyck, b. & keulen, n. 2012: searching for giant, ancient impact structures on earth: the mesoarchaean maniitsoq structure, west greenland. earth and planetary science letters 337–338, 197–210. jones, a.p., price, g.d., price, n.j., de carli, p.s. & clegg, r.a. 2002: impact induced melting and the development of large igneous provinces. earth and planetary science letters 202, 551–561. li, c., maier, w.d. & de waal, s.a. 2001: magmatic ni-cu versus pge deposits: contrasting genetic models and exploration implications. south african journal of geology 104, 309–318. lightfoot, p.c., keays, r.r., evans-lamswood, d. & wheeler, r. 2012: saturation history of nain plutonic suite mafic intrusions: origin of the voisey’s bay ni-cu-co sulfide deposit, labrador, canada. mineralium deposita 47, 23–50. maier, w.d., li, c. & de waal, s.a. 2001: why are there no major ni-cu sulfide deposits in large layered mafic–ultramafic intrusions? canadian mineralogist 39, 547–556. nielsen, b.l. 1976: economic minerals. in: escher, a. & watt, w.s. (eds): geology of greenland, 460–487. copenhagen: geological survey of greenland. secher, k. 1983: noritic rocks and associated nickel-copper-sulphide occurrences in sukkertoppen district, central west greenland. rapport grønlands geologiske undersøgelse 115, 30–34. secher, k. 2001: the pd+pt dispersion in noritic and undifferentiated mafic rocks of the archaean craton east of maniitsoq, southern west greenland. danmarks og grønlands geologiske undersøgelse rapport 2001/123, 22 pp. authors’ addresses a.a.g., t.f.k. & k.s., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: aag@geus.dk j.p., north american nickel inc., 301–260 w. esplanade, north vancouver, bc v7m 3g7, canada. i.mcd., school of earth and ocean sciences, cardiff university, p.o. box 914, cardiff cf10 3ye, uk. http://discovery.ucl.ac.uk/8800/ http://discovery.ucl.ac.uk/8800/ e2019430207-01 the landslide of 17 june 2017 in karrat fjord, central west greenland, highlighted the need for a better understanding of landslides and landslide-generated tsunamis in greenland and motivated a landslide screening project in 2018, led by the geological survey of denmark and greenland (geus; see also svennevig et al. this volume). a central part of this project was to conduct a preliminary mapping of quaternary and historical landslides in greenland – the first effort of its kind. the main objective was to establish a landslide inventory database that can be used to identify areas prone to landslides and serve as a tool for gaining a better understanding of where, when and why catastrophic landslides take place in greenland. this paper describes the workflow used to produce the preliminary landslide inventory of greenland and discusses some of the initial results. to date (june 2019), i have mapped 564 landslides with the vast majority situated in the nuussuaq basin between sigguup nunaa (svartenhuk halvø), and qeqertarsuaq (disko) in west greenland (fig. 1). the inventory mapping is mainly based on observations and analyses of remotely sensed imagery and pre-existing geological maps. the mapping coverage was not systematic for all of greenland, but focused on postglacial, potentially tsunamigenic landslides in inhabited coastal regions, i.e. on relatively large landslides on coastal slopes, mainly in west greenland and small areas of east greenland. however, smaller and inland landslides were included when they were encountered. similarly, the less inhabited parts of greenland were provisionally screened, but call for more thorough, systematic mapping in the future. existing records records of landslides in greenland are sparse and were collected ad hoc. only three large landslides are known historically, one in 1952, one in 2000 and another in 2017 (fig. 1). a previously unnamed landslide on the south coast of nuussuaq in 1952 generated a tsunami that caused one fatality, and was observed in the mining town of qullissat on northern qeqertarsuaq (dahl-jensen et al. 2004). the circumstances of the landslide are poorly known as it has not been examined in detail. this landslide is herein named the ‘1952 niiortuut landslide’ after a nearby mountain peak. the 2000 paatuut landslide (fig. 1) caused a tsunami with near-field wave heights in the order of 50 m that ravaged the then abandoned town of qullissat. it is the only landslide-tsunami event in greenland to be described in detail (pedersen et al. 2002; dahl-jensen et al. 2004). the 2017 landslide in karrat fjord was described at a reconnaissance level of detail by bessette-kirton et al. (2017) and gauthier et al. (2018), and subsequent minor landslide activity at the site was described by svennevig et al. (2019). the landslide triggered a tsunami that caused four fatalities in the nearby village of nuugaatsiaq (paris et al. 2019). additionally, kelly (1980) described a pre-historic rock avalanche deposit 25 km north-east of sisimiut at aqqutikitsoq mountain, and speculated that it occurred in the 16th or 17th century based on, among other things, the deposits relation to a local little ice age moraine. this landslide is herein referred to as the ‘aqqutikitsoq rock avalanche’. prehistoric landslide deposits were mapped in sedimentary basins and extrusive terrains on a number of published 1:100 000 scale geological maps of greenland (e.g. fig. 2a). these are mainly the map sheets covering the area from sigguup nunaa to qeqertarsuaq in west greenland (72°20´n to 69°10´n), and jameson land and the scoresby sund area in east greenland (72°00´n to 70°00´n). no landslide deposits were mapped on precambrian ‘basement’ lithologies. some of the landslide deposits on the 1:100 000 scale maps in west greenland were included on the seamless 1:500 000 scale geological map of greenland (pedersen et al. 2013). on five 1:500 000 scale quaternary geological maps in a now discontinued map series compiled by a. weidick and others (see holst et al. 2013), block/rock glaciers, mudflows, preliminary landslide mapping in greenland kristian svennevig*1 geus bulletin is free to individuals and institutions in electronic form. the author(s) retain copyright over the article contents. research article | open access geus bulletin vol 43 | e2019430207 | published online: 17 june 2019 https://doi.org/10.34194/geusb-201943-02-07 https://doi.org/10.34194/geusb-201943-02-07 e2019430207-02 and landslides were described, but inconsistently mapped. for example, landslides were only mapped in jameson land, where they are in general agreement with some of the landslides mapped on the 1:100 000 scale maps. some of the block and rock glaciers mapped by weidick and others are actually landslide deposits, as is the case with the aqqutikitsoq rock avalanche previously mentioned. three minor prehistoric landslides were recorded along the coast north of hold with hope on a 1:250 000 scale geological map of east greenland (koch & haller 1971). four relatively small prehistoric landslides around narsaq, south greenland were mapped by funder (1979). greenlandic inuit place names an unexpected data source for landslide activity indicators was the greenlandic place name database obtained from oqaasileriffik – the language secretariat of greenland. inuit place names are particularly descriptive (kleivan 1986; kruse 2012) and some of them are indicative of landslide or rock-fall activity. this includes the linguistic roots ‘nakka-’ (falling (-down)) and ‘sisuat-’ (landslide), which are directly indicative of activity (fig. 1). other words such as ‘pingu-’ (earth mound), ‘maniillat-’ (uneven/bumpy) and ‘qerrut-’ (rock pile) are indicative of geomorphological features that could be formed during landslide activity and warrant further investigation. extracting landslide data from place names is a work in progress, and several factors need to be addressed, including inconsistency issues related to local dialects and verifying that place names identified in the data200 km paatuut (2000) niiortuut (1952) aqqutikitsoq (1600 –1800) fig. 1 karrat (2017) greenland mapped landslide inuit place names that indicate: falling (‘nakka-’) landslide (‘sisuat-’) fig. 2 scoresby sund jameson land narsaq qegertarsuaq sigguup nunaq nuussuaq sisimiut hold with hope 40°w 70°n fig 1. overview map of greenland showing the 564 mapped landslides and inuit place names indicative of landslides, along with the four previously described landslides and their years of occurrence. place names mentioned in the text and the position of fig. 2 are also shown. the outline of single landslides are exaggerated for visibility on the small map scale. the concentration of landslides in the nuussuaq basin stretching from sigguup nunaa (svartenhuk halvø) to qeqertarsuaq (disko) in central west greenland, stands out. e2019430207-03 base actually match landslides in the field. until this work can be done, inuit place names that might indicate landslide activity are shown in fig. 1 but are not included in the inventory. current landslide mapping the principal data sources used for mapping were a 2 m resolution digital elevation model (dem) and the satellite images available in google earth. in areas where the resolution of the google earth images is insufficient, sentinel-2 satellite images were used (e.g. fig. 2b). in areas of particular interest such as the karrat area and the nuussuaq peninsula, the sentinel-2 images were supplemented with high-resolution oblique and nadir aerial images from the geus archive. postglacial landslides, the focus of this mapping exercise, are often clearly visible in a hillshade image created from a highresolution dem (e.g. fig. 2c). until recently, a high-resolution dem did not exist for all of greenland, but with the release of the publicly available arcticdem, greenland is now covered at a 2 m resolution (porter et al. 2018). the dem is mostly based on worldview satellite stereo images with a maximum resolution of 0.40 m. several time series of dems, some extending back to 2009, are available for the same locations, which makes it possible to follow the evolution of active landslides. in the sedimentary basins where rotational type landslides (sensu hungr et al. 2014) are most common, landslides are recognised either by the shape of their arcud a u g a a r d j e n s e n d a l c.a. niels ullanni it nerusap aqqaa ku an ne rs ui t k uu s s ua t c 2 km n 53°06’w xx 69°36’n yy ku an ne rs ui t k uu s s ua t it llanni c.a. nielse d a u g a a r d j e n s e n d a l a erusap aqqa b a fig. 2 βfph1 βf1 βi βf2 βf3 ls ls ls ls βf1u pmu fig. 2. three spatial datasets of the same area in daugaard-jensen dal, central qeqertarsuaq (disko), west greenland. see fig. 1 for location. mapped landslides are outlined by red dashed lines. a: extract from the 1:100 000-scale geological map (pedersen et al. 2001). landslide deposits (ls) are greyish purple, basalt members of the maligât formation are light green (βfph1), blue (βf1), light blue (βf1u), brown (βi), dark green (βf2) and pink (βf3). at the base of the central landslide deposit, paleocene mudstones (pmu) of the atanikerluk formation are exposed. these may have provided the failure surface for the central landslide as is also suggested for the 2000 paatuut landslide 80 km to the north (pedersen et al. 2002; dahl-jensen et al. 2004). b: sentinel-2a satellite image from 8 august 2018. c: hillshade from arcticdem (porter et al. 2018) of the area illuminated from the north-east. what appears to be a single landslide deposit in the geological map (a) is actually, on closer inspection of the arctic dem (c), two landslides: a large slide on the south-facing slope with well-defined rotated blocks, here called the daugaard-jensen dal landslide (x), and a smaller slide to the east on the north-facing slope (y). west of the daugaard-jensen dal landslide are two smaller arcuate head scarps with no landslide deposits; such deposits may be buried under the deposits of the braided river. to the east of this feature, the river has incised the toe of the daugaard-jensen dal landslide by up to 80 m. in the kuannersuit kuussuat valley to the north-west, several large landslides were also mapped. e2019430207-04 ate headscarp on otherwise smooth glacial valley sides, or by rotated blocks on the slope, which form valley-parallel ridges and grabens (fig. 2c). in general, landslide areas have a rough/hummocky morphology compared to surrounding, unfailed areas (e.g. mckean & roering 2004), and frequently have immature hydrologic drainage patterns containing many small drainless depressions. in areas underlain by crystalline rocks, landslides were not as easily recognised because headscarps were often irregular, controlled by local development of foliation and fractures. where landslide deposits were present, debris cones (molards; milana 2016; morino et al. 2019) or large boulder fields were regularly observed (e.g. in the aqqutikitsoq rock avalanche deposit). elsewhere, scarps were often the only visible sign of landslide, since the primary targets for this mapping were coastal sites where landslide deposits were located below sea level. this problem emphasises the need for detailed bathymetry in future studies. for this preliminary mapping project, submarine areas were generally not examined because of the largely poor bathymetrical data coverage. in a few locations, the bathymetrical compilation bedmachine v3 was used (morlighem et al. 2017) to identify submarine landslide deposits. once identified, landslides were mapped in arcgis as polygons with a limited number of vertices reflecting the relatively small map scale required to cover all of greenland. each polygon represents the slide perimeter from the headscarp to the toe, defined by the most distal slide deposit, if present. in some cases, only the headscarp and slide scar were mapped (e.g. fig. 2) and in a few other cases, only the landslide deposit was mapped. in the gis software, the attribute table contains data for each polygon of the mapped landslides, including a suggested or given name, references, year, month, geographical area, slope aspect, slide height (h), slide length (l), h/l ratio, local presence of bedrock dipslope (seaward dipping strata/ foliation), coincidence with landslide deposits shown in the 1:100 000 scale geological maps and generalised bedrock composition. preliminary results and outlook so far, i have mapped 564 landslides throughout greenland. the nuussuaq basin covers roughly 4.7% of the ice-free area of greenland, however, this area holds 376 of the 564 mapped landslides (67%). this confirms the previous identification of this area as a landslide ‘hot spot’ in greenland (pedersen et al. 2002; dahl-jensen et al. 2004). the geological region of karrat, within which the 2017 landslide took place, covers 3.8% of the ice free area of greenland. here, 18 landslides are mapped, comprising 3.3% of the total, and thus this region as a whole is not considered a landslide ‘hot spot’. as previously mentioned, the present mapping is preliminary and focused on the populated parts of greenland. however, all of greenland has been screened in varying detail. to produce a more accurate picture of landslide distribution throughout greenland, a more detailed and systematic mapping effort of the less populated parts of greenland is needed. furthermore, the inclusion of all available bathymetric data to map landslide deposits in the fjords would greatly add to the quality of the dataset. this is especially the case for the precambrian basement-dominated parts of west and south-east greenland, where slide scarps are often difficult to identify as they are controlled by local foliation and fracture patterns. classification of the landslides, following commonly used classifications schemes where possible (highland & bobrowsky 2008; hungr et al. 2014), is pending, and would further add to the value of the landslide inventory. additional remote sensing and fieldwork at selected sites to validate the observations would further increase the quality of the dataset. but the cost of arctic field work is a limiting factor and hence careful selection of sites is essential. another valuable add on would be sampling for absolute dates at key localities to constrain the frequencies of landslide events. available dating methods include cosmogenic nuclei analysis of headscarps and boulders, and c14-dating of sediments associated with marine landslide deposits or the small lakes formed by landslide activity (pánek 2015). a less direct method would be to examine near coastal lakes for tsunami deposits, and then date these deposits. relative age determination from remote sensing data sets using the surface roughness from high resolution dem (mccalpin 1984) or analysis of lichen cover from hyperspectral data as a proxy might be helpful in smaller regions and could be supported by a small number of absolutely dated landslides. the landslide database will be published in full with a digital appendix containing mapped polygons and attribute table when it covers all of greenland at a satisfactory scale. spatial statistical analysis on the database could be carried out in combination with available topographical, geological and climate data to improve the understanding of landslides in greenland, and in the arctic in general. this would help to identify where, when and why landslides happen in the arctic, and give some indication of the future hazards and risks from landslides in greenland. e2019430207-05 acknowledgments thanks to andrée blais-stevens and jeffrey a. coe for constructive reviews, to lotte melchior larsen for helpful comments on the manuscript and majken djurhuus poulsen for reviewing the inuit place name paragraph. references bessette-kirton, e., allstadt, k., pursley, j. & godt, j. 2017: preliminary analysis of satellite imagery and seismic observations of the nuugaatsiaq landslide and tsunami, greenland. https://www.usgs. gov/natural-hazards/landslide-hazards/science/preliminary-analysissatellite-imagery-and-seismic?qt-science_center_objects=0#qt-science_center_objects. dahl-jensen, t. et al. 2004: landslide and tsunami 21 november 2000 in paatuut, west greenland. natural hazards 31, 277–287. http://dx.doi. org/10.1023/b:nhaz.0000020264.70048.95 funder, s. 1979: the quaternary geology of the narssaq area, south greenland, grønlands geologiske undersøgelse rapport 86, copenhagen: geological survey of greenland. gauthier, d., anderson, s.a., fritz, h.m. & giachetti, t. 2018: karrat fjord (greenland) tsunamigenic landslide of 17 june 2017: initial 3d observations. landslides 15, 327-332. 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grønland. tidskriftet grønland 60, 147–156. mccalpin, j. 1984: preliminary age classification of landslides for inventory mapping. proceedings of the annual symposium on engineering geology and soil engineering 21, 99–111. mckean, j. & roering, j. 2004: objective landslide detection and surface morphology mapping using high-resolution airborne laser altimetry. geomorphology 57, 331–351. http://dx.doi.org/10.1016/ s0169-555x(03)00164-8 milana, j.p. 2016: molards and their relation to landslides involving permafrost failure. permafrost and periglacial processes 27, 271–284. http://dx.doi.org/10.1002/ppp.1878 morino, c., conway, s.j., sæmundsson, þ., kristinn, j., hillier, j., butcher, f.e.g., balme, m.r., jordan, c., argles, t., 2019. molards as an indicator of permafrost degradation and landslide processes. earth and planetary science letters 516, 136–147. https://doi.org/10.1016/j. epsl.2019.03.040 morlighem, m. et al. 2017: bedmachine v3: complete bed topography and ocean bathymetry mapping of greenland from multibeam echo sounding combined with mass conservation. geophysical research letters 44, 11051–11061. http://dx.doi.org/10.1002/2017gl074954 pánek, t. 2015: recent progress in landslide dating: a global overview. progress in physical geography 39, 168–198. http://dx.doi. org/10.1177/0309133314550671 paris, a., okal, e.a., guérin, c., heinrich, p., schindelé, f. & hébert, h. 2019: numerical modeling of the june 17, 2017 landslide and tsunami events in karrat fjord, west greenland. pure and applied geophysics, 1–23. https://doi.org/10.1007/s00024-019-02123-5 pedersen, a.k., larsen, l.m., ulff-møller, f., pedersen, g.k. & dueholm, k.s. 2001: geolgical map of greenland, 1:100 000, pingu 69 v.2 nord. copenhagen: geological survey of denmark and greenland. pedersen, m., weng, w.l., keulen, n. & kokfelt, t.f. 2013: a new seamless digital 1:500 000 scale geological map of greenland. geological survey of denmark and greenland bulletin 28, 65–68. pedersen, s.a.s., larsen, l.m., dahl-jensen, t., jepsen, h.f., krarup, g., nielsen, t., pedersen, a.k., von platen-hallermund, f. & weng, w.l. 2002: tsunami-generating rock fall and landslide on the south coast of nuussuaq , central west greenland. geology of greenland survey bulletin 191, 73–83. porter, c. et al. 2018: arcticdem v2, harvard dataverse. https://doi. org/10.7910/dvn/ohhukh svennevig, k., solgaard, a.m., dahl-jensen, t., boncori, j.p.m., larsen, t., salehi, s. & voss, p. 2019 (in press): a multidisciplinary approach to landslide monitoring in the arctic: case study of the march 2018 ml 1.9 seismic event near the karrat 2017 landslide. geological survey of denmark and greenland bulletin 43, e2019430208. how to cite svennevig, k. 2019: preliminary landslide mapping in greenland. geological survey of denmark and greenland bulletin 43, e2019430207. https://doi.org/10.34194/geusb-201943-02-07 *corresponding author: kristian svennevig | e-mail: ksv@geus.dk 1 geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350, copenhagen k, denmark. https://www.usgs.gov/natural-hazards/landslide-hazards/science/preliminary-analysis-satellite-imagery-and-seismic?qt-science_center_objects=0#qt-science_center_objects https://www.usgs.gov/natural-hazards/landslide-hazards/science/preliminary-analysis-satellite-imagery-and-seismic?qt-science_center_objects=0#qt-science_center_objects https://www.usgs.gov/natural-hazards/landslide-hazards/science/preliminary-analysis-satellite-imagery-and-seismic?qt-science_center_objects=0#qt-science_center_objects https://www.usgs.gov/natural-hazards/landslide-hazards/science/preliminary-analysis-satellite-imagery-and-seismic?qt-science_center_objects=0#qt-science_center_objects http://dx.doi.org/10.1023/b:nhaz.0000020264.70048.95 http://dx.doi.org/10.1023/b:nhaz.0000020264.70048.95 http://dx.doi.org/10.1007/s10346-017-0926-4 http://dx.doi.org/10.1007/s10346-017-0926-4 http://dx.doi.org/10.1007/s10346-013-0436-y http://dx.doi.org/10.1007/s10346-013-0436-y http://dx.doi.org/10.1016/s0169-555x(03)00164-8 http://dx.doi.org/10.1016/s0169-555x(03)00164-8 http://dx.doi.org/10.1002/ppp.1878 https://doi.org/10.1016/j.epsl.2019.03.040 https://doi.org/10.1016/j.epsl.2019.03.040 http://dx.doi.org/10.1002/2017gl074954 http://dx.doi.org/10.1177/0309133314550671 http://dx.doi.org/10.1177/0309133314550671 https://doi.org/10.1007/s00024-019-02123-5 https://doi.org/10.7910/dvn/ohhukh https://doi.org/10.7910/dvn/ohhukh https://doi.org/10.34194/geusb-201943-02-07 mailto:ksv@geus.dk geological survey of denmark and greenland bulletin 1, 555-583 555 ammonites seem to have been first reported from bornholm by malling & grönwall in 1909. they described a fauna including an ammonite which they correctly recognised as aegoceras centaurus (d’orbigny), now placed in beaniceras, from the locality of stampen, south-east of rønne (fig. 1). this established the presence of rocks dated to the pliensbachian stage of the lower jurassic marking a different zonal horizon from other ammonites reported from bornholm, and the locality has not been redescribed. malling (1914, p. 269) mentioned three species of ammonites but named only one of them, as aegoceras pettos, i.e. coeloceras pettos (quenstedt). this may be the form recorded below as coeloceras grenouillouxi (d’orbigny). malling (1920) listed six species of ammonites without description or illustration. the chief contribution to ammonite palaeontology was made by höhne (1933) who studied the clay pits at rønne where strata belonging to the hasle formation were exposed and collected ammonites in place and from the excavated material. he described and figured four species from which he concluded that the jamesoni, ibex and spinatum zones were present. this dating is revised below. however, höhne was principally concerned with the structure and petrography of the coal-bearing sediments and his palaeontological work seems to have been a sideline. lower jurassic (pliensbachian) ammonites from bornholm, baltic sea, denmark desmond t. donovan and finn surlyk the jurassic succession of the island of bornholm in the baltic sea includes the hettangian – lowermost pliensbachian rønne formation, the lower pliensbachian hasle formation, the upper pliensbachian – (?)lower aalenian sorthat formation and the (?)upper aalenian – bathonian bagå formation. ammonites are only known from the hasle formation, and all available ammonite specimens from this formation are described and figured. material reported by previous authors has been re-examined, together with previously undescribed specimens. nine genera and eleven species are recognised. the ammonites show that the rocks from which they were collected fall into the lower pliensbachian (carixian) substage. the lowermost subzone proved in the hasle formation is the basal taylori subzone of the jamesoni zone of the northwest european subboreal standard zonation, which marks a marine transgression over underlying marginal marine beds without ammonites belonging to the rønne formation. all subzones of the jamesoni zone are proved together with the valdani subzone of the ibex zone. the upper pliensbachian (domerian) substage, previously reported, is now thought to be unproved by ammonite evidence. the underlying and overlying rønne and sorthat formations are dated by dinoflagellate cysts and terrestrial palynomorphs, and the ammonite occurrences provide important control points for palynostratigraphic dating of the succession. keywords: bornholm, baltic sea, lower jurassic, ammonite fauna d.t.d., research school of geological & geophysical sciences, birkbeck college and university college london, gower street, london wc1e 6bt, uk. e-mail: ucfb03d@ucl.ac.uk f.s., geological institute, university of copenhagen, geocenter copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. geological survey of denmark and greenland bulletin 1, 555–583 (2003) © geus, 2003 556 no reports on ammonites have appeared in the literature since 1933. the collection examined here belongs to the geological museum, university of copenhagen, and includes a larger number of species than had been previously recognised, eleven species being described below. none of these are new, and although malling & grönwall (1909) and höhne (1933) both described new varieties, these are now thought to lie within the range of variation of well-known species from elsewhere. stratigraphy the jurassic of the island of bornholm in the baltic sea consists of a thickly developed hettangian–bathonian succession, whereas upper jurassic deposits are absent. four formations and a number of members are recognized (gravesen et al. 1982; koppelhus & nielsen 1994; michelsen et al. 2003, this volume). the hettangian – lowermost pliensbachian rønne formation rests on triassic redbeds of the kågeröd formation and onlaps lower palaeozoic sediments or deeply eroded crystalline basement. the rønne formation is subdivided into the lacustrine munkerup member of early hettangian age, the upper hettangian – lowermost pliensbachian paralic sose bugt member and the partly correlative upper sinemurian tidally-influenced galgeløkke member. the formation is overlain by shoreface sandstones and subordinate clay beds of the lower pliensbachian hasle formation. the age of the uppermost part of the rønne formation is somewhat uncertain. a few dinoflagellate cysts belonging to nannoceratopsis senex and mendicodinium reticulatum with known ranges from the lower pliensbachian to the middle bajocian were reported from the top of the formation by surlyk et al. (1995). the oldest ammonites from the overlying hasle formation belong to the basal pliensbachian jamesoni zone (taylori subzone), and it was tentatively suggested that the dinocysts may have an earlier appearance than previously documented (surlyk et al. 1995). two specimens of the ammonite platypleuroceras brevispina described here from the sose odde locality were probably collected in a clay ironstone bed at the top of the sose bugt member (23.1–23.4 m in surlyk et al. 1995, fig 4). they indicate the brevispina subzone of the jamesoni zone. the base of the hasle formation may thus be diachronous younging from the taylori subzone on rønne– hasle fault block hasle nyker block rønne arnager–sose fault block gudhjem trough bornholm high 10 km cretaceous lower–middle jurassic sorthat formation and middle jurassic bagå formation lower jurassic hasle formation lower jurassic rønne formation upper triassic kågeröd formation lower palaeozoic precambrian crystalline basement fault town cliff south of hasle rosmannebæk rønne lervarefabrik stampen sose odde 1 2 3 4 5 n 5 4 3 1 2 fig. 1. geological sketch map of the island of bornholm in the baltic sea showing the position of localities mentioned in the text (modified from gry 1969 and gravesen et al. 1982). the rønne–hasle block to the brevispina subzone on the arnager–sose block. the jurassic succession is capped by the paralic, lacustrine and fluvial upper pliensbachian – toarcian (?lower aalenian) sorthat formation which is overlain by the bagå formation of (?late aalenian) bajocian– bathonian age. the ammonites described in this paper are all from the heterolithic lower parts of the hasle formation. the hasle formation is 80–140 m thick and is exposed along the west coast of bornholm between the towns of rønne and hasle. a small outcrop occurs west of sose odde on the south coast (fig. 1). the formation was deposited during an overall rise in relative sea level following deposition of the paralic rønne formation (gravesen et al. 1982; surlyk & noe-nygaard 1986; koppelhus & nielsen 1994; surlyk et al. 1995; michelsen et al. 2003, this volume). the formation consists mainly of brown weathering hummocky cross-stratified finegrained sandstones. chamositic-oolitic levels and layers of pebbles and pebble-rich zones occur throughout. intercalated clay beds are common in the rønne area, where a succession belonging to the formation was formerly exposed at the locality of rønne lervarefabrik (fig. 2). in the literature and in collections, the locality is known as ‘rönne-lervarefabrik’ (höhne 1933), ‘rønne lervarefabriks ny grav’ (malling 1914), ‘rønne lervarefabrik’ (gry 1969), ‘grube rønne-lervareværk’, and ‘ny østre lergrav, rønne’. all ammonites found in the clay pits at rønne appear to come from this locality which is here called rønne lervarefabrik. höhne (1933) also described a locality which he called ‘rönne teglverk, neue grube’, but this is a different locality which exposed slightly younger coal-bearing deposits of the sorthat formation. in the section of this paper on systematic palaeontology the ammonite-bearing locality is given as rønne lervarefabrik, in each case followed by the locality names written on the original dockets in brackets. the first thorough description of a marine fauna from the jurassic of bornholm was provided by lundgren (1879) who referred it to the middle lias. his material mainly came from loose blocks and the fauna was revised by moberg (1888) in a study of a contemporaneous fauna found in situ in south-east skåne, southern sweden. a rich marine invertebrate fauna was found in 1906 at the mouth of the stampeå stream on the south coast of bornholm and was described by malling & grönwall (1909). the fauna includes 56 species and shows great resemblance to the pliensbachian of north-west germany. it was referred to the lower pliensbachian centaurus zone on the basis of the presence of the ammonite aegoceras centaurus d’orbigny var. bornholmiensis malling & grönwall. malling (1911) listed a marine fauna with 38 species of bivalves, scaphopods, gastropods, belemnites, ammonites, hybodus and plesiosaur teeth and rib bones from the low cliffs south of hasle harbour described by gravesen et al. (1982), surlyk & noe-nygaard (1986) and larsen & friis (1991). the ammonite was identified as arietites falcaries (quenstedt), i.e. arnioceras in modern nomenclature (schlegelmilch 1976, plate 20, fig. 7). malling took this to indicate the bucklandi zone (i.e. early sinemurian), which would be correct if the identification was correct. however, this is in contradiction to all the other evidence which indicates a hettangian – earliest pliensbachian age for the rønne formation and an early pliensbachian age for the base of the hasle formation. it is likely that the ammonite, which we have not seen, was misidentified. it is possible that this ammonite was a paltechioceras, as this genus ranges up to the top of the sinemurian stage. if so, it would indicate a slightly lower zonal horizon than the lowest inferred in the present paper. malling (1914) noted the presence of the ammonite aegoceras, closely related to a. maugenestii, from a locality at rosmannebæk which is only exposed at extreme low water. he also described aegoceras pettos of the jamesoni and centaurus zones which occurred together with two other ammonite species and about 30 other invertebrate species from exposures of the hasle formation at rønne lervarefabrik. the ammonites aegoceras armatum sowerby var. nodofissus quenstedt, a. caprarius quenstedt, a. maugenesti d’orbigny, coeloceras pettos quenstedt, phylloceras loscombii sowerby and oxynoticeras oppelii schloenbach were listed from the formation by malling (1920). affinities of the ammonite fauna all the genera and species now recorded from bornholm are also found in germany and england, and most of them have been recorded from north germany (hoffmann 1982). one or two species which were not listed by hoffmann (1982; e.g. coeloceras grenouillouxi) are closely related to ones that are recorded from north germany (e.g. c. pettos). the fauna is therefore typical of the boreal province. the only point worthy of remark is the absence of the genus polymorphites which is usually common in boreal jamesoni zone faunas, but in view of the small number of ammonites examined its absence may not be significant. the ammonites show that the rocks from which they were 557 558 0 5 10 15 20 25 30 35 m 70 65 60 55 50 45 40 35 fossils ammonites belemnites bivalves plants tree trunk lenticular bedding siderite cross-bedding wavy bedding structure pebbles sand clay lithology phricodoceras taylori lowermost jamesoni zone (taylori subzone) fes2 fes2 tragophylloceras cf. numismale jamesoni zone platypleuroceras caprarium jamesoni zone (brevispina subzone) f m c sandc la y si lt f m c sandc la y si lt coal fig. 2. lithological section of rønne lervarefabrik (ny østre lergrav, rønne) drawn on the basis of the written description of the stratigraphic succession in höhne (1933). collected fall into the lower pliensbachian (carixian) substage. the lowermost subzone proved in the hasle formation is the basal taylori subzone of the jamesoni zone of the nw european subboreal standard zonation (page 2003, this volume), which marks a marine transgression over underlying marginal marine beds without ammonites belonging to the rønne formation. all subzones of the jamesoni zone are proved, together with the valdani subzone of the ibex zone (fig. 3). material a total of about 35 specimens representing nine genera and eleven species are described from the hasle formation. they include tragophylloceras numismale (quenstedt), tragophylloceras sp., radstockiceras hechingense schlatter, paramicroderoceras fila (quenstedt), paramicroderoceras ? sp. indet., coeloceras grenouillouxi (d’orbigny), apoderoceras aculeatum (simpson), phricodoceras taylori (j. de c. sowerby), platypleuroceras brevispina (j. de c. sowerby), platypleuroceras submuticum (oppel), platypleuroceras caprarium (quenstedt), platypleuroceras spp. indet., uptonia lata (quenstedt), uptonia sp. indet., ?acanthopleuroceras sp. indet., beaniceras centaurus (d’orbigny), and ammonites gen. et sp. indet. all specimens are from the collection of the geological museum, university of copenhagen, and no new material has been added as the clay pits were abandoned many years ago. the material is housed in the geological museum, university of copenhagen (mguh 25595–25627). systematic palaeontology full synonymies for some species would be very long; references are only given here to citations relevant to the discussion and to occurrences on bornholm. suborder phylloceratina arkell 1950 family juraphyllitidae arkell 1950 genus tragophylloceras hyatt 1900 tragophylloceras numismale (quenstedt 1845) plate 1, figs 1–6 1845 ammonites heterophyllus numismalis quenstedt, p. 100, plate 6, figs 4a, b, 5a, b, non figs 3a, b, 5c. 1885 ammonites heterophyllus numismalis quenstedt, p. 291, plate 37, figs 8–11, 21. ?1920 phylloceras loscombii sowerby – malling, p. 55. 1933 phylloceras heterophyllum numismale quenstedt – höhne, p. 56, plate 14, figs 2, 3. 1964 tragophylloceras numismale (quenstedt) – howarth & donovan, p. 295, plate 48, fig. 5. 1976 tragophylloceras numismale (pompeckj) – schlegelmilch, plate 2, figs 1, 2. 559 stage subzonezone northwest european province zonule (nw european s.s.) figulinum capricornus maculatum luridum valdani ibex davoei masseanum masseanum jamesoni jamesoni jamesoni brevispina brevispina/brevispinoides polymorphus polymorphus taylori taylori figulinum angulatum crescens/samontaensis capricornus lataecosta maculatum sparsicosta luridum crassum rotundum alisiense centaurus venarense actaeon valdani maugenesti arietiforme pettos tenuilobos/submuticum biruga nodogigas lo w er p lie ns ba ch ia n fig. 3. pliensbachian zonal and subzonal subdivision (modified from page 2003, this volume). on the basis of the ammonite evidence, all four of the subzones of the jamesoni zone and the valdani subzone of the ibex zone (shaded subzones) are represented on bornholm. 1980 tragophylloceras numismale (quenstedt) – schlatter, plate 1, fig. 1a, b. type specimen. the original of quenstedt (1845, plate 6, fig. 5a, b) was designated lectotype by buckman (1912, p. viii). discussion. full synonymy and description can be found in howarth & donovan (1964, p. 295). stratigraphical horizon. howarth & donovan (1964) noted that in well-dated north-west european localities, the species is restricted to the lower part of the jamesoni zone. material. a specimen now numbered 1987/94 was described and figured by höhne (1933) from rønne lervarefabrik (grube rønne-lervareværk), and is now refigured (plate 1, figs 3–6). it is 67 mm in diameter and wholly septate, lacking, as höhne remarked, the body chamber and part of the phragmocone. specimens numbered 1976/804[2, 3] (plate 1, figs 1, 2) are identified as tragophylloceras cf. numismale. they are parts of phragmocones, the larger being about 110 mm in diameter. tragophylloceras sp. fig. 4a 1933 phylloceras heterophyllum numismale quenstedt – höhne, p. 56, plate 14, fig. 1. höhne (1933) described and figured from rønne lervarefabrik (grube rønne-lervareværk) a fragment of a smooth ammonite showing parts of typical tragophylloceras suture lines. it does not appear to be among the specimens examined by us. this is apparently the phylloceras listed from bed 11 of his section (höhne 1933, p.11). it could be t. numismale as höhne thought, but is not now considered to be specifically identifiable. suborder ammonitina zittel 1884 superfamily psilocerataceae hyatt 1867 family oxynoticeratidae hyatt 1875 genus radstockiceras buckman 1918 for synonymy of the genus, see donovan et al. (1981, p. 137). radstockiceras hechingense schlatter 1980 plate 1, figs 7–9; plate 2, figs 1, 2 1854 ammonites oxynotus numismalis oppel, p. 84, plate 2, fig. 10a–c. 1856 ammonites oxynotus numismalis oppel – quenstedt, p. 119, plate 14, fig. 1. 1885 ammonites oxynotus numismalis oppel – quenstedt, p. 289, plate 37, figs 1–7. 1920 oxynoticeras oppelii schloenbach – malling, p. 55. 1925 metoxynoticeras numismale (quenstedt) – spath, p. 112, footnote 4. 1962 metoxynoticeras complanosum (simpson) – howarth, p. 105, plate 15, fig. 4a, b. 1976 radstockiceras complanosum (simpson) – schlegelmilch, plate 23, fig. 3 (original of quenstedt 1885, plate 37, fig. 1, refigured). 1980 metoxynoticeras numismale (oppel) – schlatter, p. 50, plate 1, figs 4, 5. 1980 radstockiceras hechingense schlatter, p. 52. type. the original of oppel (1854, plate 2, fig. 10a–c) is the holotype. it is no longer extant (schlatter 1980, p. 51). spath (1925) wrongly stated the original of quenstedt (1885, plate 37, fig. 4) to be the lectotype. nomenclature. schlatter (1980) pointed out that ammonites oxynotus numismalis oppel (1854) was, when published, a homonym of ammonites heterophyllus numismalis quenstedt (1845). he therefore proposed the new specific name hechingense for oppel’s (1854) species. simpson’s am. complanatus (1843) was replaced on account of homonomy by his am. complanosus (1855). the holotype of the latter species (whitby museum: no. 239) was first figured by howarth (1962, plate 15, fig. 4a, b) but is now said to be lost (schlatter 1980, p. 51). the bornholm examples agree well with oppel’s (1854) somewhat stylised drawing of the lost holotype and with schlatter’s (1980) photographic figures. the species is strongly compressed, the whorl thickness being about 25% of the diameter. the venter is fastigate rather than sharp as in some oxycones. on the body chamber of large individuals, the fastigate venter is lost, the whorl becomes thicker and the venter broadly rounded as shown by quenstedt (1885, plate 37, fig. 3), a specimen with whorl height 110 mm and thickness 47 mm. the ribbing persists to varying diameters on different individuals. the more nearly complete example from bornholm (plate 1, figs 7–9) is about 107 mm in diameter, retaining some of the body chamber, but much larger specimens are known from elsewhere. 560 561 stratigraphical horizon. the genus radstockiceras probably ranges from the upper raricostatum zone to the mid-davoei zone (donovan 1994). in the type pliensbachian in southern germany, it is almost restricted to the polymorphus subzone (schlatter 1980, p. 52; table 2) but in northern germany it seems to range through much of the jamesoni zone and possibly into the ibex zone (hoffmann 1982, p. 152). it is not, therefore, a good indicator of stratigraphical horizon. material. nos 1976/804[1], 1976/80, both from rønne lervarefabrik (ny østre lergrav, rønne). family eoderoceratidae spath 1929 genus paramicroderoceras dommergues et al. 1994 it was previously noted by donovan (1990, p. 258–259) that a number of evolute, bituberculate species have been commonly, but wrongly, referred to microderoceras by continental authors. donovan (1990) then placed them in tetraspidoceras spath 1926, as did spath himself (e.g. spath 1938, p. 5, fig. 1a). however, tetraspidoceras as defined by the type species, am. quadrarmatus dumortier, has inner whorls with the umbilicus only about 35% of the diameter, ornamented with numerous fine ribs and pairs of small tubercles at intervals. the evolute, strongly tuberculate form is only developed on the body chamber. in contrast, the forms now under discussion show little, if any, change of shell form or ornament with growth. dommergues et al. (1994, p. 36) have also pointed out that tetraspidoceras is inappropriate for these species and have proposed the new genus paramicroderoceras for them. paramicroderoceras fila (quenstedt 1884) plate 2, figs 10–12 1884 ammonites armatus fila quenstedt, p. 205, plate 25, fig. 7; plate 26, figs 2, 3. a c1 d1 d2 e1 e2 c2 bfig. 4. ammonites reproduced from höhne (1933, plate 14, figs 1, 3–9); all are shown at natural size. the material has not been seen by us and is probably lost. a, tragophylloceras sp. rønne lervarefabrik. b, tragophylloceras numismale (quenstedt). rønne lervarefabrik. c1, 2; d1, 2, ammonite gen. et sp. indet. rønne lervarefabrik. e1, 2, ammonite gen. et. sp. indet. rønne lervarefabrik. 1976 microderoceras fila (quenstedt) – schlegelmilch, plate 24, fig. 4 (quenstedt 1884, plate 25, fig. 7 refigured). 1980 microderoceras fila (quenstedt) – schlatter, p. 54, plate 2, fig. 2. nomenclature. the status of quenstedt’s ‘trinomina’ is doubtful. hölder (1958, p. 22) has stated that this name had not been previously used. as it has been employed as a species name by authors it is retained here. type. the original of quenstedt (1884, plate 25, fig. 7) was refigured by schlegelmilch (1976, plate 24, fig. 4) and stated to be lectotype. it is in the geological collections at the university of tübingen, germany. discussion. the single example from bornholm agrees well with the lectotype. it is a phragmocone 100 mm in diameter. the last preserved part of the shell has been damaged by injury during life. the figured examples attributed to this species are all phragmocones, the later whorls and body chamber being unknown. am. armatus bimacula quenstedt (1884, p. 207, plate 26, fig. 4; a ‘neotype’ was figured by schlegelmilch 1976, plate 24, fig. 3) differs little from am. arm. fila except by greater whorl thickness. the two names are probably to be considered as synonyms. stratigraphical horizon. schlatter (1980) reported the species from the jamesoni zone, near the junction of the taylori and polymorphus subzones, and hoffmann (1982, p. 156, plate 12, figs 1, 2a, b) recorded similar forms as probably from the lower part of the jamesoni zone. such a horizon agrees well enough with other bornholm material, but the genus is long-ranging and therefore not a very good stratigraphical indicator. material. one specimen, 1987/93, from rønne lervarefabrik (ny østre lergrav, rønne). paramicroderoceras ? sp. indet. plate 2, fig. 13 ?1920 aegoceras armatum sowerby var. nodofissus quenstedt – malling, p. 55. a single specimen, 1976/806, from ny østre lergrav, rønne, is the internal mould of a fragment of a septate whorl. it has a compressed, flat-sided whorl section, and strong ribs, possibly bituberculate, each pair separated by three or four weaker ribs. it bears some resemblance to hyperderoceras planarmatus (quenstedt) in schlegelmilch (1976, plate 27, fig. 2). that species is the type species of parahyperderoceras schlatter (1980, p. 62) which schlatter regarded as a subgenus of hyperderoceras spath, but which dommergues (1987, p. 99) regarded as a synonym of epideroceras (coeloderoceras). in view of the fragmentary nature of the present specimen further discussion seems unhelpful. stratigraphical horizon. schlatter (1980, plate 3, fig. 7) illustrated a somewhat similar form, with stronger and coarser ornament, from the late taylori/earliest polymorphus subzone. material. 1976/806 from rønne lervarefabrik (ny østre lergrav, rønne). family coeloceratidae haug 1910 genus coeloceras hyatt 1867 coeloceras grenouillouxi (d’orbigny 1844) plate 2, figs 6–9 1844 ammonites grenouillouxi d’orbigny, p. 307, plate 96, figs 1–6. 1914 aegoceras pettos malling, p. 269. 1920 aegoceras pettos quenstedt – malling, p. 55. 1980 coeloceras grenouillouxi (d’orbigny) – schlatter, p. 73, plate 6, figs 1–3. 1994 coeloceras grenouillouxi (d’orbigny) – r. mouterde and j.-l. dommergues in fischer, p. 85, plate 22, figs 4 (lectotype), 5a, b, 6, 7a, b. types. there are 15 surviving syntypes in the d’orbigny collection of which one was designated lectotype, and figured, by schlatter (1980, p. 73, plate 6, fig. 1). the lectotype and three other syntypes were also figured by r. mouterde and j.-l. dommergues (in: fischer 1994). discussion. the single specimen from bornholm agrees well with the lectotype. schlatter (1980), studying the german material from the type area of c. pettos, treated c. grenouillouxi as separate from c. pettos. however, dommergues & mouterde (1978, p. 346), reviewing abundant material from the type locality of c. grenouillouxi at cottards (cher), france, concluded that the assemblage showed considerable variability and they regarded c. grenouillouxi as a synonym of c. pettos. 562 r. mouterde and j.-l. dommergues (in: fischer 1994, p. 85) also thought that the two ‘species’ “pourraient n’être en réalité que deux morphes de la même espèce”. dommergues (1994) has recently claimed to recognise dimorphism in the genus. stratigraphical horizon. in southern germany, southern england and burgundy, coeloceras is found in the upper part of the jamesoni subzone. dommergues (1987, p. 25) recognised a horizon of evolutus (pettos) and page (2003, this volume) a pettos zonule. it is likely that this horizon is present in bornholm. material. no. 1987/83 from rønne lervarefabrik (ny østre lergrav, rønne) (1987/92 is the impression of 1987/83). this may be the specimen recorded as aegoceras pettos by malling (1914) from rønne lervarefabrik (described by him as rønne lervarefabriks ny grav, his locality 5; fig. 1). genus apoderoceras buckman 1921 apoderoceras aculeatum (simpson 1843) plate 2, figs 3–5 1843 ammonites aculeatus simpson, p. 27. 1843 ammonites marshallani simpson, p. 24. 1855 ammonites aculeatus simpson, p. 66. 1855 ammonites marshallani simpson, p. 62. 1876 aegoceras aculeatus (simpson) – blake, p. 278, plate 7, fig. 4. 1880 aegoceras leckenbyi wright, plate 30, figs 1–7. 1882 aegoceras leckenbyi wright, p. 344. 1884 ammonites aculeatus simpson, p. 99. 1884 ammonites marshallani simpson, p. 94. 1913 apoderoceras aculeatum (simpson) – buckman, plate 72a–c. 1954 apoderoceras aculeatum (simpson) – donovan, p. 35. 1962 apoderoceras aculeatum (simpson) – howarth, p. 109, plate 15, fig. 5a, b. nomenclature. simpson’s (1843) unillustrated specimen was first figured by blake (1876), and later, photographically, by buckman (1913). ammonites marshallani was not figured until howarth (1962) who regarded it as a synonym of aculeatum. ammonites decussatum simpson (1843, p. 25) is probably also a synonym (howarth 1962, p. 109). the species is somewhat variable and several other specific names of authors are probably synonyms. type. the holotype of apoderoceras aculeatum is lost. a paratype (whitby museum: no. 177) was figured by buckman (1913, plate 72a–c). discussion. the species when complete is about 300 mm in diameter. the coeloceras-like innermost whorls are succeeded by a finely ribbed or striate, tuberculate, stage which gives way to a body chamber ornamented with stout ribs bearing ventro-lateral tubercles, (e.g. hoffmann 1982, plate 37). the single specimen from bornholm is a typical example of the inner whorls, only 79 mm in diameter. stratigraphical horizon. in northern europe apoderoceras occurs in the taylori subzone. the present species probably comes from this subzone. material. m 1933/68, from rønne lervarefabrik (ny østre lergrav, rønne). family phricodoceratidae spath 1938 genus phricodoceras hyatt 1900 phricodoceras taylori (j. de c. sowerby 1826) plate 4, figs 9–13; plate 6, figs 3–5 1826 ammonites taylori j. de c. sowerby, p. 23, plate 514, fig. 1. 1843 ammonites cornutus simpson, p. 31. 1855 ammonites cornutus simpson, p. 71. 1884 ammonites taylori sowerby – simpson, p. 105. 1911 phricodoceras cornutum simpson – buckman, plate 32. 1933 aegoceras taylori sowerby var. bornholmiensis höhne, p. 52, plate 12, figs 1–4, 6, 7. 1961 phricodoceras taylori (j. de c. sowerby) – dean et al., plate 68, fig. 5a, b. 1976 phricodoceras cornutum (simpson) – schlegelmilch, plate 28, fig. 1? types. sowerby’s (1826) figured specimen and presumed holotype of ammonites taylori is lost. it was from a boulder in glacial till at happisburgh, norfolk, england. there are four syntypes of höhne’s (1933) var. bornholmiensis from rønne lervarefabrik (rønne lervare563 fabriks ny grav). it has not been possible to trace the originals. they were probably lost from the collections at the university of greifswald during world war two. the replicas seen by us are numbered as follows: 1976/809 = original of höhne 1933, plate 12, figs 1, 3 1976/810 = original of höhne 1933, plate 12, fig. 2 1976/811 = original of höhne 1933, plate 12, figs 6, 7 1976/812 = original of höhne 1933, plate 12, fig. 4 note that some of höhne’s illustrations have been trimmed, and that the figures are not half natural size as stated on the plate. the replicas are figured here as plate 4, figs 9–13 and plate 6, figs 3–5. the original of höhne’s exemplar 3, i.e. of his plate 12, figs 6, 7 (replica no. 1976/811), is here designated the lectotype of aegoceras taylori sowerby var. bornholmiensis höhne. it is refigured here (plate 6, figs 3–5). nomenclature. there are two common forms of small phricodoceras, one of which has strong tubercles at the ventral ends of the ribs; tubercles half-way along the ribs, if present, die out at a small size (c. 20 mm). this is p. taylori. ammonites cornutus simpson 1843 is a synonym, as simpson himself recognised in 1884 (p. 105). the other form has the inner tubercles persisting to a larger size, and this is ammonites quadricornutus simpson 1855 (holotype figured by buckman 1911, plate 32). unfortunately, schlegelmilch in his picturebook of lias ammonites has confused the two forms and illustrates p. quadricornutum under the name of p. taylori (schlegelmilch 1976, plate 27, fig. 3) and p. taylori as p. cornutum (schlegelmilch 1976, plate 28, fig. 1). phricodoceras costatum (quenstedt) as figured by schlatter (1980, p. 77, plate 6, fig. 4) appears close to p. taylori. schlatter (1980, p. 78) refers the bituberculate species from pliensbach to p. nodosum (quenstedt). however, am. taylori nodosus (quenstedt 1846, p. 136) was, when proposed, a junior homonym of am. nodosa de roissy and at least two other prior uses. it appears to the present writers that quenstedt’s nodosus can be accommodated in simpson’s quadricornutus. aegoceras taylori sowerby var. bornholmiensis of höhne (1933) is here considered a synonym of p. taylori. discussion. phricodoceras includes both small species (taylori, quadricornutum) and large ones (lamellosum d’orbigny, non j. de c. sowerby) and is probably dimorphic, as noted by dommergues (1978) who interpreted ammonites lamellosus d’orbigny as the macroconch of p. taylori. only the small forms (microconchs) have been reported from bornholm. they are typical examples of p. taylori. stratigraphical horizon. in tethyan areas, phricodoceras has a long stratigraphical range, from the upper sinemurian to the upper pliensbachian. in northwest europe, however, it characterises the lowermost part of the jamesoni zone (taylori subzone), and it is likely to mark this horizon in bornholm. in burgundy (dommergues 1987) and in southern germany (schlatter 1980), phricodoceras occurs mainly above apoderoceras (taylori and nodogigas zonules of page 2003, this volume), and it is possible that both these levels are represented in bornholm. material. 1976/809–812, replicas of höhne’s (1933) four specimens referred to above, from bed 53 in the section at rønne lervarefabrik recorded by höhne (1933, p. 11, 12); 1987/91 from rønne lervarefabrik (ny østre lergrav, rønne). family polymorphitidae haug 1887 genus platypleuroceras hyatt 1867 platypleuroceras brevispina (j. de c. sowerby 1827) plate 3, figs 1–6 1827 ammonites brevispina j. de c. sowerby, p. 106, plate 556, fig. 1 (fig. 2 cited in the text in error). 1845 ammonites natrix rotundus (var. α) quenstedt, p. 85, plate 4, fig. 17a–c. 1845 ammonites natrix oblongus (var. ß) quenstedt, p. 85, plate 4, fig. 16a–c. 1880 aegoceras brevispina (j. de c. sowerby) – wright, plate 32, figs 2, 3. 1882 aegoceras brevispina (j. de c. sowerby) – wright, p. 361. 1885 ammonites amplinatrix quenstedt, p. 257, plate 32, fig. 7. 1925 platypleuroceras brevispinoides tutcher & trueman, p. 649, plate 40, fig. 2. 1954 platypleuroceras brevispina (j. de c. sowerby) – donovan, p. 36. 1961 platypleuroceras brevispina (j. de c. sowerby) – dean et al., plate 69, fig. 1. 564 1976 platypleuroceras brevispina (j. de c. sowerby) – schlegelmilch, plate 29, figs 3, 4. 1980 platypleuroceras brevispina (sowerby 1827) – schlatter, p. 95, plate 8, fig. 1 (with synonymy). 1980 platypleuroceras oblongum (quenstedt 1845) – schlatter, p. 98, plate 8, fig. 2; plate 9, fig. 1 (with synonymy). 1980 platypleuroceras rotundum (quenstedt 1845) – schlatter, p. 101, plate 9, figs 2, 3 (with synonymy). type. j. de c. sowerby’s only specimen, regarded as the holotype, was refigured by wright (1880, plate 32, figs 2, 3), dean et al. (1961, plate 69, fig. 1) and by schlegelmilch (1976, plate 29, fig. 3). it is in the british museum (natural history) no. 43915. details of the types of ammonites natrix oblongus, ammonites natrix rotundus and platypleuroceras brevispinoides are given by schlatter (1980). after studying abundant material, dommergues (1987, p. 141) concluded that the other names included in the above synonymy had been applied to forms which fall within the range of variation of p. brevispina. the specimens from bornholm are typical of the species. dommergues (1987, p. 141, 142) believed the species to be dimorphic. the bornholm examples are all macroconchs according to dommergues’ interpretation, complete with body chambers at diameters of about 100 mm. stratigraphical horizon. the species characterises the brevispina subzone of the jamesoni zone, and ranges throughout the subzone according to schlatter (1980, table 2). in burgundy, dommergues (1987, p. 29) regarded the species as characteristic of the lower part of the subzone, replaced by species of the group of p. submuticum in the upper part. material. nos 1886/224 and 1976/803, both from sose odde; 1976/702 from rønne lervarefabrik (ny østre lergrav, rønne). platypleuroceras submuticum (oppel 1856) plate 4, figs 1–6; plate 6, figs 8–10 1854 ammonites natrix oblongus quenstedt – oppel, p. 35, plate 1, fig. 5a, b. 1856 ammonites submuticus oppel, p. 278. 1980 platypleuroceras submuticum (oppel) – schlatter, p. 103, plate 9, fig. 5; plate 10, figs 1, 2 (with synonymy). 1980 platypleuroceras enzenharti schlatter, p. 107, plate 11, figs 1, 2 (holotype). 1987 platypleuroceras submuticum (oppel) – dommergues, p. 142, plate 9, figs 1–13. type. the holotype is the ammonite figured by oppel (1854) as quenstedt’s am. natrix oblongus and renamed by him in 1856. it is in the bayerische staatssammlung at münchen no. as viii 156 and was refigured by schlatter (1980, plate 9, fig. 5). the holotype of platypleuroceras enzenharti is in the staatliches museum für naturkunde, stuttgart, no. 23567. dommergues (1987) has discussed the characters and variability of the species. he regarded p. enzenharti schlatter as a variant of p. submuticum. compared with p. brevispina the species is typically compressed with flat-sided whorls and closer ribbing. the inner row of spines is absent. the examples from bornholm are typical. by analogy with p. brevispina they are presumably macroconchs. no. 1975/1491 (plate 6, figs 8–10) is still septate at a diameter of about 120 mm. stratigraphical horizon. the species is found in the upper part of the brevispina subzone (schlatter 1980, table 2; dommergues 1987, p. 29). material. nos 1975/1491, 1987/94, both from rønne lervarefabrik (rønne lervarefabriks ny grav). no. 1976/808 from the same locality, an impression seen in side view only, may belong to the species. platypleuroceras caprarium (quenstedt 1856) plate 3, figs 7–9; plate 5, figs 1–21 1856 ammonites caprarius quenstedt, p. 131, plate 16, fig. 1. 1884 ammonites caprarius quenstedt, p. 243, plate 30, figs 37–42. 1920 aegoceras caprarius (quenstedt) – malling, p. 55. 1933 aegoceras m.f. valdani-maugenesti höhne, p. 54, plate 13, figs 3?, 4, 5. 1976 platypleuroceras caprarium (quenstedt) – schlegelmilch, p. 63, plate 29, fig. 5 (neotype). 1977 platypleuroceras caprarium (quenstedt) – schlatter, p. 10, plate 2, fig. 1a, b. 1980 polymorphites caprarius (quenstedt) – schlatter, p. 92. 1982 platypleuroceras caprarium (quenstedt) – hoffmann, p. 194, plate 20, figs 1a, b, 2a, b, 3a, b, 4, 5a–d, 6. 565 1987 ‘platypleuroceras’ caprarium (quenstedt) – dommergues, p. 152, plate 6, figs 45–48; plate 7, figs 1–4. type. the specimen figured by quenstedt (1856) is lost. a neotype was designated and figured by schlegelmilch (1976). discussion. the species is characterised by compressed whorls, close, regular ribbing, and strong ventral ornament. there are small but equally prominent inner and outer spines. the bornholm examples are typical of the species. plate 13, fig. 3 in höhne (1933) is the side view only of a specimen which we have not seen. it may belong to p. caprarium. bornholm individuals with body chamber are 55–60 mm in diameter when complete; specimen no. 1987/86 is about 61 mm in diameter with three-quarters of a whorl of body chamber, and an unnumbered specimen (plate 5, figs 7–9) has a body chamber of about three-quarters of a whorl at 58 mm diameter and is complete with the aperture of the shell. the species is smaller than later species of platypleuroceras. dommergues (1987, p. 152) regarded p. caprarium as a macroconch and illustrated examples ranging from about 45 mm to 84 mm in diameter. he illustrated as p. (m) sp. 3 (dommergues 1987, plate 6, figs 23–30) specimens ranging from 29 to 44 mm diameter which he thought might be the microconchs of the species. the bornholm examples would all be macroconchs according to dommergues’ (1987) definition. stratigraphical horizon. detailed collecting of the type pliensbachian by schlatter (1980) showed the species to be the earliest species of platypleuroceras, though occurring wholly above the range of phricodoceras. hoffmann (1982, p. 260) recorded it below, as well as in association with, phricodoceras in north germany. material. nos 1987/86–90 from rønne lervarefabrik (ny østre lergrav, rønne). the specimen 1987/87 is a replica of that figured by höhne (1933, plate 13, figs 4, 5). note that höhne’s figs 4 and 5 are 0.7 and 0.86 respectively of true size, not 0.5 as stated. no. 1831 (2 complete specimens and a body chamber fragment) from rønne lervarefabrik (rønne lervarefabriks nye grav). no. 2021 from the same locality (?). two unnumbered complete specimens. platypleuroceras spp. indet. plate 4, fig. 7; plate 6, fig. 6 1933 aegoceras m.f. valdani-maugenesti d’orbigny – höhne, p. 54, plate 13, figs 1, 2. höhne (1933) illustrated in side view only two small platypleuroceras, the originals of which have not been seen by us. the material examined includes a (body chamber?) fragment, 1976/807, from rønne lervarefabrik (ny østre lergrav, rønne), and an impression, 1976/813 (plate 4, fig. 7), found loose on the beach north of rønne. none of these examples is considered to be identifiable at species level. genus uptonia buckman 1897 uptonia lata (quenstedt) (sensu schlatter 1980) plate 6, figs 1, 2 1980 uptonia ex gr. lata (quenstedt) – schlatter, p. 114, plate 13, fig. 1 (with synonymy). 1980 uptonia confusa (quenstedt) – schlatter, p. 114, plate 13, figs 2a, b, 3a, b; plate 14, fig. 1a, b (with synonymy). 1987 uptonia lata (quenstedt) (sensu schlatter) – dommergues, p. 133, plate 10, figs 1–5. type. quenstedt’s (1845) figured specimen is not known to exist. nomenclature. quenstedt (1845, p. 88; 1849, plate 4, fig. 1) described and figured ammonites jamesoni var. latus. his illustration is of a fragment of an outer whorl which cannot be satisfactorily interpreted. dommergues (1987) has adopted the name in the interpretation of schlatter (1980). the species is distinguished from u. jamesoni (j. de c. sowerby), as interpreted by the neotype designated by donovan & forsey (1973, p. 12, plate 4, fig. 3a, b), by having fewer ribs throughout, and stronger ornament on the body chamber. dommergues (1987, p. 115, fig. 28) regarded the species as later than u. jamesoni in his evolutionary lineage. the species is large. schlatter (1980, plate 14, fig. 1) figured (as u. aff. confusa) a wholly septate example 117 mm in diameter, and dommergues (1987, plate 10, figs 1, 2) figured one still septate at 132 mm diameter. the bornholm specimen is smaller, with about threeeighths of a whorl of body chamber at a diameter of 566 about 120 mm, and inner whorls mostly preserved as an impression. stratigraphical horizon. in the type pliensbachian the species occurs throughout the jamesoni subzone and just ranges into the succeeding ibex zone (schlatter 1980). material. no. 1987/85, recorded only as from bornholm island. uptonia sp. indet. plate 6, fig. 7 a septate whorl fragment from rønne lervarefabrik (ny østre lergrav, rønne), no. 1976/801, is not considered to be specifically identifiable. family acanthopleuroceratidae ?acanthopleuroceras sp. indet. plate 4, fig. 8 ?1920 aegoceras maugenesti d’orbigny – malling, p. 55. an incompletely preserved ammonite, no. 1987/84, differs from the species of platypleuroceras and uptonia recorded above in having nearly smooth inner whorls, followed by coarse, blunt ribs. it bears some resemblance to forms figured, for example by schlatter (1980, plate 16, fig. 3), as acanthopleuroceras maugenesti (d’orbigny). if this identification is correct it indicates the ibex zone, valdani subzone. the locality is recorded only as bornholm island. family liparoceratidae genus beaniceras s.s. buckman 1913 beaniceras centaurus (d’orbigny 1844) 1844 ammonites centaurus d’orbigny, p. 266, plate 76, figs 3–6. 1909 aegoceras centaurus d’orbigny var. bornholmiensis malling & grönwall, p. 288, plate 11, figs 9–11. 1918 beaniceras senile buckman, plate 126. 1919 beaniceras centaurus (d’orbigny) – buckman, plate 146. 1924 beaniceras senile buckman, plate 126a. 1938 beaniceras centaurus (d’orbigny) – spath, p. 107, plate 16, fig. 9a, b. 1938 beaniceras centaurus (d’orbigny) var. cherensis spath, p. 108, plate 6, fig. 7a–c. 1938 beaniceras centaurus (d’orbigny) var. elegans spath, p. 108, plate 6, fig. 8a, b; plate 16, fig. 9c, d. 1938 beaniceras centaurus (d’orbigny) var. subcrassa spath, p. 108, plate 16, figs 8a, b. 1938 beaniceras senile buckman – spath, p. 109, plate 15, figs 2a, b. 1978 beaniceras centaurus (d’orbigny) – dommergues & mouterde, plate 3, figs 11a–c, 12a–c, 13a–c. 1985 androgynoceras (subgen. nov.) centaurus (d’orbigny) – phelps, plate 1, figs 14, 15a, b. 1994 beaniceras centaurus (d’orbigny) – j.-l. dommergues and r. mouterde in fischer, p. 69, plate 23, figs 7a–c, 8a–c, 9a, b (syntypes figured). types. there are twenty specimens of am. centaurus in the d’orbigny collection (j.-l. dommergues and r. mouterde in: fischer 1994, p. 69). no lectotype has been designated. the holotype of beaniceras senile (original of buckman 1918, plate 126) is in the british geological survey collection no. 47092. aegoceras centaurus d’orbigny var. bornholmiensis malling & grönwall was evidently based on the single figured specimen which is therefore the holotype. we have not seen this specimen. discussion. phelps (1985) regarded the species as belonging to a new (but unnamed) subgenus (of androgynoceras). pending the naming of such a subgenus, the species is here retained in beaniceras. d’orbigny (1844) illustrated only the cadicone inner whorls. spath (1938, p. 109) separated b. senile buckman from b. centaurus on the basis of its contracted bodychamber with arched venter in contrast to the flatter venter of the preceding whorl. however, material from the type area figured by dommergues & mouterde (1978) shows that in the adult, the body chamber, about half a whorl long and 20 mm diameter, ceases to expand with growth and corresponds with the form named b. senile by buckman (1918). phelps (1985), who revised the stratigraphy and taxonomy of beaniceras, figured (plate 1, fig. 14) as centaurus an ammonite which is indistinguishable from the holotype of b. senile, which he presumably regarded as a synonym. malling & grönwall (1909) described and figured as aegoceras centaurus d’orbigny var. bornholmiensis an ammonite from their locality at stampen, bornholm 567 which is very similar to beaniceras senile, and is therefore placed in synonymy. it should be noted, however, that if it is desired to uphold buckman’s (1924) species b. senile, then the name bornholmiensis has priority. stratigraphical horizon. phelps (1985) recognised a centaurus zonule near the top of the valdani subzone, and this has been adopted by page (2003, this volume). material. malling and grönwall’s (1909) figured specimen has not been seen by us, and its present location is not known. ammonite gen. et sp. indet. fig. 4c1–2, d1–2 1933 amaltheus cf. spinatus bruguière – höhne, p. 55, plate 14, figs 4–7. discussion. the whorl fragments were identified by höhne (1933) as close to amaltheus spinatus, i.e. pleuroceras in modern nomenclature. if correct this would indicate the domerian stage, spinatum zone. however, the fragments do not look like pleuroceras as far as can be seen from the poor illustration, reproduced here (fig. 4). they appear to show bifurcating ribs, which do not occur in pleuroceras, and do not show the interrupted keel characteristic of that genus. the v-shaped ribs arise from an umbilical tubercle, with thickened, forwardly curved outer ends. it is to be noted that while the specimen illustrated as fig. 4 in plate 14 (höhne 1933) is shown with its anterior end upwards, the corresponding ventral view (fig. 5) is wrongly oriented, with the anterior end downwards. there are several jurassic and cretaceous ammonite genera which have the characters described above, and it is not considered safe to identify these fragments from the figures. material. two whorl fragments figured by höhne (1933), from rønne lervarefabrik (rønne lervareverk). stated to be in the bornholm museum at rønne. this material has not been seen by us. ammonite gen. et sp. indet. fig. 4e1–e2 1933 amaltheus cf. spinatus bruguière – höhne, p. 55, plate 14, figs 8, 9. material. a whorl fragment found by höhne in bed 3 of rønne lervarefabrik (höhne 1933, p. 11); 66.4–69.4 m in fig. 2. discussion. this whorl fragment was also wrongly identified by höhne (1933) as amaltheus. it has straight ribs each of which appears to bear three tubercles or spines. it is not considered to be identifiable from the figures. acknowledgements we thank walter kegel christensen for assistance in locating some of the specimens, jan aagaard for photography, merete vesterager for word processing, christian hagen for drafting, and kevin n. page and lars h. nielsen for reviewing the manuscript. references blake, j.f. 1876: cephalopoda. in: tate, r. & blake, j.f. (eds): the yorkshire lias, 261–330. london: john van voorst. buckman, s.s. 1909–1930: yorkshire type ammonites (1, 2) and type ammonites (3–7), 790 plates. london: wheldon & wesley. specific volumes/parts cited in this paper: buckman, s.s. 1911: yorkshire type ammonites 1 (parts 3–5). buckman, s.s. 1912: yorkshire type ammonites 1 (parts 6–8). buckman, s.s. 1913: yorkshire type ammonites 2 (parts 9–11). buckman, s.s. 1918: yorkshire type ammonites 2 (parts 15–17). buckman, s.s. 1919: yorkshire type ammonites 2 (part 18); type ammonites 3 (parts 19, 20). buckman, s.s. 1924: type ammonites 5 (parts 44–49). dean, w.t., donovan, d.t. & howarth, m.k. 1961: the liassic ammonite zones and subzones of the north-west european province. bulletin of the british museum (natural history), geology 4, 435–505. dommergues, j.-l. 1978: un cas de dimorphisme sexuel chez une ammonite carixienne phricodoceras taylori (j. de c. sowerby 1826) eoderoceratidae spath 1929. bulletin scientifique de bourgogne 31, 41–45. dommergues, j.-l. 1987: l’évolution chez les ammonitina du lias moyen (carixien, domerien basal) en europe occidentale. documents des laboratoires de géologie de facúlte des sciences de lyon 98, 272 pp. dommergues, j.-l. 1994: the jurassic ammonite coeloceras: an atypical example of dimorphic progenesis elucidated by cladistics. lethaia 27, 143–152. dommergues, j.-l. & mouterde, r. 1978: les faunes d’ammonites du carixien inférieur et moyen du gisement des cottards (cher). géobios 11, 345–365. 568 dommergues, j.-l., ferretti, a. & meister, c. 1994: les faunes d’ammonites du sinémurien de l’apennin central (marches et toscane, italie). bollettino della società paleontologica italiana 33, 13–42. donovan, d.t. 1954: synoptic supplement to t. wright’s ‘monograph on the lias ammonites of the british islands’ (1878–86), 54 pp. london: palaeontographical society. donovan, d.t. 1990: sinemurian and pliensbachian ammonite faunas of central italy. in: pallini, g. et al. (eds): atti ii convegno internazionale: ‘fossili, evoluzione, ambienti’, pergola 1987, 253–262. donovan, d.t. 1994: evolution in some early jurassic ammonites: asteroceratinae, oxynoticeratidae and related forms. in: pallini, g. (ed.): proceedings of the 3rd pergola international symposium ‘fossili, evoluzione, ambienti’, palaeopelagos special publication 1, 383–396. rome: università ‘la sapienza’. donovan, d.t. & forsey, g.f. 1973: systematics of lower liassic ammonitina. university of kansas paleontological contributions paper 64, 18 pp. donovan, d.t., callomon, j.h. & howarth, m.k. 1981: classification of the jurassic ammonitina. in: house, m.r. & senior, j.r. (eds): the ammonoidea. systematics association special volume 18, 101–155. d’orbigny, a. 1844: paléontologie française. terrains jurassiques. i. céphalopodes, 193–312. paris: masson. fischer, j.-c. (ed.) 1994: révision critique de la paléontologie française d’alcide d’orbigny. volume i. céphalopodes jurassiques, 340 pp + facsimile of d’orbigny 1842–1851. paris: masson. gravesen, p., rolle, f. & surlyk, f. 1982: lithostratigraphy and sedimentary evolution of the triassic, jurassic and lower cretaceous of bornholm, denmark. danmarks geologiske undersøgelse serie b 7, 51 pp. gry, h. 1969: megaspores from the jurassic of the island of bornholm. meddelelser fra dansk geologisk forening 19, 69–89. hoffmann, k. 1982: die stratigraphie, paläogeographie und ammonitenführung des unter-pliensbachium (carixium, lias gamma) in nordwest-deutschland. geologisches jahrbuch reihe a 55, 442 pp. höhne, r. 1933: beiträge zur stratigraphie, tektonik und paläogeographie des südbaltischen rhät-lias, insbesondere auf bornholm. abhandlungen aus dem geologisch-paläontologischen institut der universität greifswald 12, 1–105. hölder, h. 1958: vorschläge für die behandlung von f.a. quenstedt’s nomenklatur. paläontologische zeitschrift 32, 18–23. howarth, m.k. 1962: the yorkshire type ammonites and nautiloids of young and bird, phillips, and martin simpson. palaeontology 5, 93–136. howarth, m.k. & donovan, d.t. 1964: ammonites of the liassic family juraphyllitidae in britain. palaeontology 7, 286–305. koppelhus, e.b. & nielsen, l.h. 1994: palynostratigraphy and palaeoenvironments of the lower to middle jurassic bagå formation of bornholm, denmark. palynology 18, 139–194. larsen, o.h. & friis, h. 1991: petrography, diagenesis and porewater evolution of a shallow marine sandstone (hasle formation, lower jurassic, bornholm, denmark). sedimentary geology 72, 269–284. lundgren, b. 1879: bidrag til kännedomen om juraformationen på bornholm. festskrift till universitet i köpenhamn vid dess jubileum 1879, 29 pp. lund: lunds universitet. malling, c. 1911: hasle-sandstenens alder. meddelelser fra dansk geologisk forening 3, 629–631. malling, c. 1914: de jespersenske buelag i lias paa bornholm. meddelelser fra dansk geologisk forening 4, 265–270. malling, c. 1920: den marine lias og wealden-aflejringer paa bornholm. meddelelser fra dansk geologisk forening 5, 55–57. malling, c. & grönwall, k.a. 1909: en fauna i bornholms lias. meddelelser fra dansk geologisk forening 3, 271–316. michelsen, o., nielsen, l.h., johannessen, p.n., andsbjerg, j. & surlyk, f. 2003: jurassic lithostratigraphy and stratigraphic development onshore and offshore denmark. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 147–216 (this volume). moberg, j.c. 1888: om lias i sydöstra skåne. sveriges geologiske undersökning. serie c 99, 86 pp. oppel, a. 1854: der mittlere lias schwabens. jahresheft des vereins für vaterländische naturkunde in württemberg 10, 39–136. [usually cited as 1853; hoffmann (1982, p. 351) however, stated that it was printed in 1854]. oppel, a. 1856–1858: die juraformation englands, frankreichs und des südwestlichen deutschlands, nach ihren einzelnen gliedern eingetheilt und verglichen. jahresheft des vereins für vaterländische naturkunde in württemberg 12–14, 857 pp. stuttgart: ebner & seubert. [oppel’s work appeared in issues 2, 3 of volume 12, issues 2, 3 of volume 13 and issues 2, 3 of volume 14, with the corresponding page numbers 1856: p. 1–438; 1857: p. 439–694; 1858: p. 695–857]. page, k.n. 2003: the lower jurassic of europe: its subdivision and correlation. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 23–59 (this volume). phelps, m.c. 1985: a refined ammonite biostratigraphy for the middle and upper carixian (ibex and davoei zones, lower jurassic) in north-west europe and stratigraphical details of the carixian–domerian boundary. geobios 18, 321–362. quenstedt, f.a. 1845–1849: petrefactenkunde deutschlands. i. die cephalopoden, 580 pp. tübingen: fues. quenstedt, f.a. 1856: der jura [first part], 576 pp. tübingen: laupp. quenstedt, f.a. 1884: die ammoniten des schwäbischen jura. i. der schwarze jura, 97–240. stuttgart: schweizerbart. quenstedt, f.a. 1885: die ammoniten des schwäbischen jura. i. der schwarze jura, 241–440. stuttgart: schweizerbart. schlatter, r. 1977: the biostratigraphy of the lower pliensbachian at the type locality (pliensbach, württemberg, sw germany). stuttgarter beiträge zur naturkunde serie b (geologie und paläontologie) 27, 29 pp. schlatter, r. 1980: biostratigraphie und ammonitenfauna des unter-pliensbachium im typusgebiet (pliensbach, holzmaden und nürtingen; württemberg, sw-deutschland). stuttgarter 569 570 beiträge zur naturkunde serie b (geologie und paläontologie) 65, 261 pp. schlegelmilch, r. 1976: die ammoniten des süddeutschen lias, 212 pp. stuttgart: gustav fischer verlag. simpson, m. 1843: a monograph of the ammonites of the yorkshire lias, 60 pp. london: simpson, marshall. simpson, m. 1855: the fossils of the yorkshire lias; described from nature, 149 pp. london: whittaker. simpson, m. 1884: the fossils of the yorkshire lias; described from nature, 2nd edition, 256 pp. whitby, london: wheldon. sowerby, j. de c. 1826: the mineral conchology of great britain 6 (part 88). london: j. de c. sowerby. sowerby, j. de c. 1827: the mineral conchology of great britain 6 (part 95). london: j. de c. sowerby. spath, l.f. 1925: notes on yorkshire ammonites. the naturalist 1925 819, 107–112. spath, l.f. 1938: a catalogue of the ammonites of the liassic family liparoceratidae in the british museum (natural history), 191 pp. london: trustees of the british museum. surlyk, f. & noe-nygaard, n. 1986: hummocky cross-stratification from the lower jurassic hasle formation of bornholm, denmark. sedimentary geology 46, 259–273. surlyk, f., arndorff, l., hamann, n.e., hamberg, l., johannessen, p.n., koppelhus, e.b., nielsen, l.h., noe-nygaard, n., pedersen, g.k. & petersen, h.i. 1995: high-resolution sequence stratigraphy of a hettangian–sinemurian paralic succession, bornholm, denmark. sedimentology 42, 323–354. tutcher, j.w. & trueman, a.e. 1925: the liassic rocks of the radstock district (somerset). quarterly journal of the geological society of london 81, 595–666. wright, t. 1880: monograph on the lias ammonites of the british islands, part 3, 165–264. london: palaeontographical society. wright, t. 1882: monograph on the lias ammonites of the british islands, part 5, 329–400. london: palaeontographical society. manuscript received 17 march 2000; revision accepted 31 may 2000. 571 plates 1–6 572 plate 1 all figures are at natural size. figs 1, 2. tragophylloceras cf. numismale (quenstedt). no. 1976/804[2], from rønne lervarefabrik (ny østre lergrav, rønne). mguh 25595. figs 3–6. tragophylloceras numismale (quenstedt). no. 1987/94, from rønne lervarefabrik (grube rønne-lervareverk), previously figured by höhne (1933, plate 14, figs 2, 3). mguh 25596. figs 7–9. radstockiceras hechingense schlatter. no. 1976/805, from rønne lervarefabrik (ny østre lergrav, rønne). mguh 25597. 573 1 4 6 7 8 9 2 3 5 574 plate 2 all figures are at natural size. figs 1, 2. radstockiceras hechingense schlatter. no. 1976/804[1], from rønne lervarefabrik (ny østre lergrav, rønne). mguh 25598. figs 3–5. apoderoceras aculeatum (simpson). no. m 1933/68, from rønne lervarefabrik (ny østre lergrav, rønne). mguh 25599–25600. figs 6–9. coeloceras grenouillouxi (d’orbigny). no. 1987/92, rønne lervarefabrik (new quarry of the rønne lervarefabrik). fig. 9 is specimen no. 1987/83, an impression of the umbilicus of 1987/92. mguh 25601–25602. figs 10–12. paramicroderoceras fila (quenstedt). no. 1987/93, from rønne lervarefabrik (ny østre lergrav, rønne). mguh 25603. fig. 13. paramicroderoceras ? sp. indet. no. 1976/806, from rønne lervarefabrik (ny østre lergrav, rønne). mguh 25604. 575 1 2 5 10 11 12 9 13 6 7 8 3 4 576 plate 3 all figures are at natural size. figs 1–6. platypleuroceras brevispina (j. de c. sowerby). 1, no. 1976/803; 2, 3, no. 1886/224, both from sose odde; 4–6, no. 1976/702, from rønne lervarefabrik (ny østre lergrav, rønne). mguh 25605–25607. figs 7–9. platypleuroceras caprarium (quenstedt), inner whorls. no. 1987/90, from rønne lervarefabrik (ny østre lergrav, rønne). mguh 25608. 577 1 2 4 5 6 3 7 8 9 578 plate 4 all figures are at natural size. figs 1–6. platypleuroceras submuticum (oppel). 1–4, no. 1987/94, 5, 6, no. 1987/808, both from rønne lervarefabrik (rønne lervarefabriks ny grav). mguh 25609–25610. fig. 7. platypleuroceras sp. indet. no. 1976/813, found loose on the beach, north of rønne. mguh 25611. fig. 8. ?acanthopleuroceras sp. indet. no. 1987/84. bornholm. mguh 25612. figs 9–13. phricodoceras taylori (j. de c. sowerby). 9, 10, no. 1976/810; 11, no. 1976/812; 12, 13, no. 1987/91, from rønne lervarefabrik (ny østre lergrav, rønne). mguh 25613–25615. 579 1 2 3 4 5 6 8 7 9 10 11 12 13 580 plate 5 all figures are at natural size. figs 1–21. platypleuroceras caprarium (quenstedt). 1, 2, no. 2021, rønne lervarefabrik (rønne lervarefabriks nye grav); 3, 4, no number, rønne lervarefabrik (rønne lervarefabriks nye grav), with aperture preserved; 5, 6, no. 1831 (one of three so numbered), rønne lervarefabrik (rønne lervarefabriks nye grav); 7–9, no number; 10, 11, no. 1987/87, previously figured by höhne (1933, plate 13, figs. 4, 5); 12–14, no. 1987/88; 15–17, no. 1987/89; 18–21, no. 1987/86 (2 specimens); nos 10–21 from rønne lervarefabrik (ny østre lergrav, rønne). mguh 25616–25622. 581 1 8 13 12 14 16 21 2018 19 15 17 7 9 10 11 2 3 4 6 5 582 plate 6 all figures are at natural size. figs 1, 2. uptonia lata (quenstedt) (sensu schlatter). no. 1987/85, exact locality not known. mguh 25623. figs 3–5. phricodoceras taylori (j. de c. sowerby). no. 1976/811, from rønne lervarefabrik (ny østre lergrav, rønne). mguh 25624. fig. 6. platypleuroceras sp. indet. no. 1976/807, from rønne lervarefabrik (ny østre lergrav, rønne). mguh 25625. fig. 7. uptonia sp. indet. no. 1976/801, from rønne lervarefabrik (ny østre lergrav, rønne). mguh 25626. figs 8–10. platypleuroceras submuticum (oppel). no. 1975/1491, from rønne lervarefabrik (rønne lervarefabriks ny grav). mguh 25627. 583 1 2 3 4 5 6 7 9 8 10 geological survey of denmark and greenland bulletin 7, 2004, p 69-72 69 previous investigations by the geological survey of denmark and greenland (geus) and exploration companies have demonstrated that some of the kimberlites in west greenland are diamond bearing, making the region an important target for diamond prospecting. high-resolution hyperspectral (hs) remote sensing data have been successfully used for the location of kimberlitic rocks, e.g. in australia and africa. however, its potential as a viable method for the mapping of kimberlite occurrences in arctic glaciated terrain with high relief was previously unknown. in july–august 2002, geus conducted an airborne hyperspectral survey in central west greenland (fig. 1) using the commercially available hymap hyperspectral scanner operated by hyvista corporation, australia. data were processed in 2003, and in 2004 follow-up field work was carried out in the kangerlussuaq region to test possible kimberlites indicated by the hs data (fig. 1). the project was financed by the bureau of minerals and petroleum, government of greenland. hyperspectral data and field work the hymap airborne hyperspectral scanner (cocks et al. 1998), developed by integrated spectronics, sydney, australia, delivers high accuracy, calibrated radiance data over 126 channels covering the wavelength range between 400 and 2500 nm with 15–20 nm bandwidth. the hymap system also generates the flight line ephemeris data (x, y, z and aircraft attitude data) utilising its satellite navigation system (dgps) and integrated inertial monitoring unit (imu). these data are necessary for georectification and advanced processing of the hs image data. the survey area in central west greenland was flown with the following specifications: data coverage 7500 km2 number of lines 54 line kilometres 3500 nominal pixel size 4 metres overlap per line 20% approximate ground speed 140 knots (280 km/h) detection of kimberlitic rocks in west greenland using airborne hyperspectral data: the hypergreen 2002 project tapani tukiainen and leif thorning geological survey of denmark and greenland bulletin 7, 69–72 (2005) © geus, 2005 granitic intrusions (s.l.) archaean gneiss reworked in the palaeoproterozoic supracrustal rocks orthogneisses calc-alkaline intrusions arfersiorfik and sisimiut suites thrustt t tt t t t t t t t t t t ttt t kangerlussuaq 50 km 66°n greenland 51°w sisimiut ka ng er lus su aq sukkertoppen iskappe inland ice greenland fig. 1. simplified geological map of the study region in west greenland. the coverage of the hypergreen 2002 survey is indicated by the black frame. the red frame outlines the map area of fig. 3. 70 at the same time, a field programme was carried out to measure a number of spectra from selected kimberlite occurrences to establish the spectral characteristics of the kimberlitic rocks and their erosion products in west greenland (tukiainen et al. 2003). spectral basis for the mapping of kimberlitic rocks kimberlites consist of predominantly ultramafic material that has crystallised in situ, and commonly host megacrysts formed in the upper mantle from the kimberlite magma and mantle derived xenoliths (dunite, lherzolite, wehrlite, harzburgite, eclogite and granulite) incorporated during magma transport. common matrix minerals are olivine, phlogopite, perovskite, spinel, chromite, diopside, monticellite, apatite, calcite and fe-rich serpentine. the most interesting minerals with respect to hyperspectral mapping are phlogopite, fe-rich serpentine (antigorite) and calcite; these minerals have characteristic spectral responses in the short wave infrared (swir) spectral region (2.0–2.5 µm). comparison of the hymap spectrum of kimberlite to the spectra measured with a field instrument at the same locality demonstrates a close match (fig. 2). hymap data analysis atmospheric correction to fully exploit the possibilities of hyperspectral image data delivered ‘at sensor radiance data’, they must be converted to surface reflectance data. the small size of potential targets and the relatively subtle spectral characteristics as established by the ground truth survey, demonstrated that the rugged terrain conditions of west greenland require the use of atmospheric correction methods, which take sensor viewing geometry and terrain information into consideration. the conversion of the data to surface reflectance was done using the atcor-4 package (richter & schläpfer 2002). the photogrammetric laboratory at geus produced a detailed digital elevation model which was used as terrain information for the atcor-4 system. spectral mapping the field measurements have shown that the spectral response from kimberlitic rocks within wavelengths of 2.0–2.5 µm is remarkably uniform. thus the simplest way to locate the kimberlitic rocks is to use selected characteristic kimberlite field spectra as end members for the spectral processing. the spectral angle mapper (sam, kruse et al. 1993) was used in this project for comparing the hs image spectra to selected, characteristic kimberlite field spectra. the algorithm determines the similarity between two spectra by calculating the ‘spectral angle’ between them, treating them as vectors in space with dimensionality equal to the number of bands. the method is not sensitive to the unknown gain factor and all possible illuminations are treated equally. this is an important advantage when processing data acquired in the hypergreen 2002 project, where illumination levels vary between the flight routes and even within a single flight line. the sam algorithm calculates the angular distance (in radians) between each spectrum in the image and the refer2000 antigorite (usgs) v al u e (s p ec tr a o ff se t fo r cl ar it y) kimberlite (hymap) kimberlite (fs) phlogopite (usgs) weathered kimberlite (fs) 2100 2200 wavelength nm 2300 2400 fig. 2. comparison of laboratory mineral spectra (usgs, clark et al. 1993) to the kimberlite spectra measured by field instruments (fs) and airborne hyperspectral scanner (hymap, locality k12 in figs 3, 4). the field instrument covers the range 0.4–2.5 µm at a higher spectral resolution. note that hymap, though lower resolution, resolves the key spectral features near 2.3 µm. the spectral features in hymap and field spectra of kimberlite near 2.3 µm are distinctly subdued when compared to the laboratory spectra of phlogopite and antigorite. this is probably due to the fact that the kimberlite spectrum is a linear mixture of all materials occurring within the hymap pixel. ence spectra. the ‘rule’ image for each end member shows the actual distance between each spectrum in the image and the reference spectrum. low values of angular distance denote high similarities between the spectra. mapping results a subset of the area to which the hs mapping was applied in 2003 (tukiainen & krebs 2004) is here used to illustrate the use of hs data for the mapping of kimberlite occurrences (fig. 3). the area was chosen because reliable field follow-up information is available. the largest known exposed kimberlite occurrence (locality k12; figs 3, 4) where the exposure correspond to 4–5 hymap image pixels, was readily detected by the sam method, even when mapping is based only on the phlogopite mineral spectra measured in laboratory conditions. the limited field follow-up resulted in discovery of a number of kimberlite exposures and boulder floats. the newly discovered kimberlite occurrences are typically small, outcrops rarely exceeding the nominal pixel size of 4 × 4 m (fig. 4). known limitations of the method and sources of error and misclassifications the hymap hyperspectral scanner is an optical sensor and can only detect targets which are visible. illumination conditions caused by a combination of high and complex topography imply that parts of the terrain are in shadow where the poor signal/noise ratio camouflages the subtle spectral features. the high atmospheric water vapour content, typically above and adjacent to major fjords and nearby valleys, suppress the signal from the short-wave infrared part of the spectrum thereby increasing the noise level of the image data. extreme illumination conditions (areas adjacent to snow/ice and bright surfaces) and complex, steep topography may also create image-processing artefacts. the applied hs mapping strategy is based on detection of the minerals phlogopite, serpentine and calcite when these are present as rock forming minerals. these minerals, or combinations of them, also commonly occur in rock types other than kimberlite (ultramafic rocks, and various carbonate rocks, carbonate-veined shear zones, altered and weathered mafic and ultramafic rocks). the field follow-up in 2004 showed that the most common source of error was caused by weathered and altered exposures of kangâmiut dykes, which are the most common mafic rocks in the survey area. the spectral characteristics of the kangâmiut dyke rocks were studied in more detail in 2004, and the processing scheme was hereafter adjusted to better distinguish them from kimberlite outcrops. conclusion the airborne hyperspectral data acquired by the hymap hyperspectral sensor are capable of detecting kimberlite occurrences in west greenland when the exposed surface of kimberlite outcrops and/or the weathering products approaches or exceeds the image pixel size (4–5 m). the sometimes unfavourable terrain and illumination conditions may, however, seriously affect the detection success rate. the success rate for detecting rocks with phlogopite, serpentine and carbonates as main constituents is good, although distinction between rock types is more problematic. the rugged terrain conditions of west greenland and the small size of the potential targets, typically corresponding to less than one or a few image pixels, and the relatively subtle spectral characteristics near 2.3 µm in the swir spectrum, require the use of atmospheric correction methods which take the sensor viewing geometry and terrain information into consideration. 71 kangerlussuaq k1 k12 k10 n3 n fig. 3. results from the kimberlite mapping from an area covering a part of the surveyed area (location indicated in fig. 1). known and discovered kimberlite occurrences (in situ occurrences and boulder floats) are shown on the map; those detected by hyperspectral mapping are shown with circles. 72 references clark, r.n., swayze, g.a., gallagher, a., king, t.v.v. & calvin, w.m. 1993: the u.s. geological survey, digital spectral library: version 1: 0.2 to 3.0 µm. united states geological survey open file report 93-592, 1326 pp. cocks, t., jenssen, a., stewart, i., wilson, i. & shields, t. 1998: the hymap airborne hyperspectral sensor: the system, calibration, and performance. in: schaepman, m., schläpfer, d. & itten, k.i. (eds): proceedings of the 1st earsel workshop on imaging spectroscopy, 37–43. zurich, switzerland, 6–8 october, 1998. paris: european association of remote sensing laboratories. kruse, f.a., lefkoff, a.b., boardman, j.b., heidebrecht, k.b., shapiro, a.t., barloon, p.j. & goetz, a.f.h. 1993: the spectral image processing system (sips). interactive visualisation and analysis of imaging spectrometer data. remote sensing of environment 44, 145–163. richter, r. & schläpfer, d. 2002: geo-atmospheric processing of airborne imaging spectrometry data. part 2: atmospheric/topographic correction. international journal of remote sensing 23, 2631–2649. tukiainen, t. & krebs, j.d. 2004: mineral resources of the precambrian shield of central west greenland (66° to 70°15´n). part 4. mapping of kimberlitic rocks in west greenland using airborne hyperspectral data. danmarks og grønlands geologiske undersøgelse rapport 2004/45, 39 pp. + 1 dvd. tukiainen, t., krebs, j.d., kuosmanen, v., laitinen, j. & schäffer, u. 2003: field and laboratory reflectance spectra of kimberlitic rocks, 0.35–2.5 µm, west greenland. danmarks og grønlands geologiske undersøgelse rapport 2003/43, 25 pp. a b c k k k k d fig. 4. kimberlite outcrop localities k1, k10, k12 and n3 indicated in fig. 3. a: strongly weathered kimberlite (k1). b: typical small exposure and weathered material (k10). c: hanging wall of a kimberlite (k) dyke (n3). d: largest known exposure of kimberlite in west greenland (outlined in red), measuring 30 × 5 m (k12). authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: tt@geus.dk geological survey of denmark and greenland bulletin 38, 2017, 29-32 29 geothermal resources in the deep subsurface in many parts of denmark have the potential to form a central component in the future danish energy supply for district heating. geothermal energy is sustainable and environmentally friendly and independent of climatic and seasonal variations, in contrast to solar and wind energy. furthermore, geothermal plants may be integrated with other green energy supplies. the sandstone reservoirs from which the warm geothermal water is extracted may also act as temporary storage for excess heat e.g. from industrial production processes or from solar-heated water in summer periods when the demand for heating is low. therefore, there are many good reasons to include geothermal energy in denmark’s energy mix. despite this, only three geothermal plants exist at present at thisted, copenhagen and sønderborg (fig. 1). several district heating companies have, however, shown interest in geothermal energy and have taken the first step towards estimating if suitable geological conditions are present within their respective district heating areas. this has been done by analysing geological and geophysical data gathered from the nearest deep wells and seismic surveys, in some cases supplemented with new seismic data. although these analyses generally show promising geothermal potential, hesitation prevails when it comes to drilling the actual geothermal wells. deep drilling is complicated and expensive, but necessary in order to deduce if productive reservoir sandstones are present that can produce the required volumes of geothermal water. in order to mitigate the geological risks and facilitate utilisation of geothermal energy, publicly supported initiatives financed by research grants have been undertaken for the last 40 years. thereby our knowledge of the danish subsurface and the reservoir properties of deep geothermal sandstones has considerably increased, and fundamental uncertainties regarding subsurface structures and resources have been reduced. the many promising results are publicly available via the newly established geothermal webgis portal at the geological survey of denmark and greenland (vosgerau et al. 2016). furthermore, the industry is now taking more interest in geothermal exploration and sees it as a promising business case into which it is willing to invest and share the risks associated with expensive wells. the public sector facilitates this development by supporting research projects via grants from the energy technology development and demonstration programme of the danish energy agency (eudp) and the innovation fund denmark. these projects involve research institutes, district heating companies, private companies and other stakeholders. the present paper deals with the outcome of one of these projects called the geothermal pilot hole, financially supported by the eudp. the project elucidates e.g. how drilling can be made less expensive by focusing on geothermal sandstone reservoirs at depths shallower than c. 2200 m, thereby allowing the use of small rigs suitable towards a geothermal exploration well in the gassum formation in copenhagen henrik vosgerau, ulrik gregersen, lars kristensen, sofie lindström, anders mathiesen, carsten m. nielsen, mette olivarius and lars henrik nielsen 100 km denmarkdenmark t s sweden petrophysical log data no data or poor quality reasonable or good seismic data coverage and quality very poor poor resonable good very good 12°e 56°n lavø-1 karlebo-1/1a slagelse-1 stenlille-1 stenlille-19 sjælland amager fig. 2 20 km• • • •• • • • margretheholm-1/1a /-2 margretheholm-1/1a /-2 fig. 1. coverage and quality of seismic and petrophysical log data from deep wells in sjælland. the quality indeces reflect to which degree the data can be used to extract information about major geothermal sandstone reservoirs in the deep subsurface. the margretheholm wells are part of the existing geothermal plant in copenhagen. stars mark the approximate positions of the two areas of interest. on the inset map the locations of the thisted (t) and sønderborg (s) plants are shown. © 2017 geus. geological survey of denmark and greenland bulletin 38, 29–32. open access: www.geus.dk/publications/bull 3030 for operation in urban areas. for comparison, the existing geothermal plant in copenhagen utilises water from the lower triassic bunter sandstone formation at a depth of c. 2.6 km. another important part of the project is to provide a well-constrained prognosis of relevant reservoir parameters (depth, thickness, transmissivity, production capacity, temperature, etc.) of geothermal sandstone reservoirs of the upper triassic – lower jurassic gassum formation within two prospect areas of special interest in northern copenhagen (fig. 1). this activity will provide the necessary background to select the location of an exploration well in a future phase of the project yet to be granted. previous studies have shown that the subsurface of copenhagen contains large quantities of geothermal energy which may form a substantial contribution to domestic heating for hundreds of years to come, and the two areas in northern copenhagen (stars, fig. 1) have already been selected as relevant, based on suitable geological conditions and infrastructure. copenhagen is a major city with a substantial demand for heating and like other danish cities it has a well-established district heating network, and for these reasons it is an obvious site for geothermal energy. the gassum formation constitutes the most well-known sandstone reservoir in denmark and is exploited for geothermal energy in thisted and sønderborg and for gas storage at stenlille. it is dominated by fine to medium-grained sandstones alternating with darker-coloured claystones, siltstones and thin coal seams. the sand was deposited in the danish basin mainly as marine shoreface sand in relatively continuous and widely distributed bodies, as well as deposits in river channels, estuaries and lagoons. in the copenhagen area, the gassum formation occurs in depths of around 2000 m and has a temperature of c. 60°c (balling et al. 2016), sufficiently high to make a district heating plant economically profitable. furthermore, the depth is shallow enough to prevent serious diagenetic alteration under high pressure and temperature conditions which might reduce the porosity and permeability of the reservoir sandstones (kristensen et al. 2016). the results of the project phases conducted so far illustrates e.g. that the subsurface geological conditions may vary considerably within a city area, thus influencing the geothermal potential. geological database the critical subsurface geological information from deep wells and seismic lines in central and north-eastern sjælland (fig. 1) controls the reservoir prognosis for the two prospect areas in northern copenhagen. the seismic coverage is reasonable around these two areas, especially because an e–w-trending seismic line of very good quality occurs immediately north of them. detailed analysis of this and other nearby lines has been used to identify and estimate the depth and thickness of the gassum formation in the two prospect areas. the two areas are separated by the nnw–sse-striking amager fault which forms part of a major regional fault zone, along which the easternmost part of sjælland has been down-faulted (fig. 2). the seismic mapping reveals that the gassum formation is thicker and occurs at a deeper level in the eastern area than in the western area. deep wells in north-eastern sjælland are limited to margretheholm-1/1a and -2, karlebo-1/1a and lavø-1 from which no cores of the gassum formation exist, and n e am ager fault zone base gassum fm base gassum fm base chalk western prognosis area w m ar gr et he ho lm -1 m ar gr et he ho lm -2 am ager fault zone gassum fm base chalk twt (ms) -1100 -1400 -1300 -1200 eastern prognosis area 2 km seismic line hgs001 seismic line hgs1e11 seismic line hgs8n13 sea level øresund sweden fjerritsle v fm fjerritslev fm gassum fm fjerritslev fm base chalk gassum fm view position and direction 5 km n am ager fault zone fig. 2. seismic lines in 3d view seen towards northern copenhagen and margretheholm in northern amager. the view position and direction are shown on the inserted location map, as is the overall trend of the amager fault zone. the coloured surface reflects the depth and morphology of the base of the gassum formation, clearly illustrating the different, fault-controlled positions of the two prospect areas west and east of the fault. the thickness of the gassum formation, marked on the seismic profiles, increases considerably from west to east across the amager fault zone. depths are shown as seismic two-way travel time, twt. 31 the petrophysical log data from lavø-1 are of poor quality. however, a large amount of good quality petrophysical log and core data exist from the stenlille area in the central part of sjælland, c. 60 km west of the two areas of interest (fig. 1), although not all of the well data are equally relevant for setting up a reservoir prognosis for the gassum formation in the prospect areas. representative well data have been chosen based on e.g. sequence stratigraphic and biostratigraphic analysis and on similarities to the prospect areas in terms of structural setting, depositional environment, inferred distance from the palaeoshoreline, sediment sources, petrography, burial depth, diagenetic alternation, etc. in general, the data from the margretheholm and karlebo-1/1a wells are considered the most relevant for the eastern prospect area, as also these wells are located on the eastern, down-faulted side of the amager fault (fig. 2). the locations of the lavø-1 and the stenlille wells west of the amager fault justify a higher weighting of data from these wells in the prognosticating of the western area. estimated reservoir values based on the seismic data, the gassum formation is estimated to be around 200 m thick with its top at c. 2000 m below sea level in the eastern of the project areas in copenhagen, and around 150 m thick with its top at c. 1750 m below sea level in the western area (table 1). core and petrophysical well log data form the basis for estimating the reservoir properties of the sandstones including local porosity–permeability models (fig. 3). the porosity and permeability values are estimated to be slightly lower in the eastern than in the western area, most likely related to its deeper burial depth and corresponding higher pressure–temperature conditions (table 1). the reservoir transmissivity, given by multiplying the estimated thickness of potential reservoir sand with the estimated reservoir permeability, is an important parameter as it expresses the overall performance of the reservoir. as a rule of thumb, the transmissivity of a sandstone interval in the danish subsurface should be greater than 10 darcy-m in order to constitute a potential geothermal reservoir. both areas fulfil this criterion as the estimated reservoir transmissivity for the gassum formation is 25 and 28 darcy-m, respectively, for the two areas. although the estimated porosity, permeability and transmissivity values are slightly higher in the western area, this does not necessary qualify this as better for geothermal exploitation. this is because the geothermal water of the gassum formation in the eastern area benefits from being hotter than in the western area (65°c versus 57°c, table 1) as a consequence of its greater depth. reservoir model simulations reservoir simulations in both of the prospect areas have shed further light on the suitability of the gassum formation for geothermal exploitation. the reservoir data and interpreted regional seismic surfaces have thus been used to simulate flow rates and the time span before cooled water from injection wells will reach the production wells. in each of the simulations, separate production and injection wells supplemented with a vertical spud well were used. simulation runs with different well spacings show that the distance between the production and injection wells at depth can be kept as low as 900 m without cold-water breakthrough at the production well within the simulation period of 25 years. given that the injection and production wells would typically originate from the same surface table 1. estimated reservoir values for the gassum formation in two potential areas§ prospect areas in copenhagen western eastern macro reservoir parameters depth to top of formation (m below sea level) 1750 2000 thickness (m) 150 200 thickness, potential reservoir sand (m) 75 80 water-conducting properties (reservoir sand) porosity (%) 25 21 gas permeability (md) 300 250 reservoir permeability (md) 375 313 reservoir transmissivity (kh) (dm) 28 25 temperature (°c), middle of formation 57 65 §see kristensen et al. (2016) for details. mah-2 (swc) stenlille-1 to -6 (ccal) copenhagen model stenlille model copenhagen: y = 0.000377x4.3622 stenlille: y = 0.000222x4.3622 10 20 30 10 100 1000 10000 pe rm ea bi lit y (m d ) porosity % 0 1 fig. 3. porosity–permeability models based on gas measurements on core material at laboratory conditions. the ‘copenhagen model’ represents the eastern prospect area and is based on measurements on sidewall cores from the margretheholm-2 well, whereas the ‘stenlille model’ represents the western prognosis area and is based on conventional core analysis data from several stenlille wells. the methodology for establishing local porosity–permeability models is described in kristensen et al. (2016). 3232 position, short distances between the injection and production wells at the reservoir level are preferred in order to minimise the inclination of the well trajectory. this will lower the drilling risks, as drilling generally becomes more complicated with increasing inclination. overall, the simulations showed suitable production capacities for both locations but that the eastern location is more favourable because of higher production and injection rates for the same pressure applied to the wells, a more favourable production temperature profile, as well as thicker reservoir intervals which will delay breakthrough of cold water from the injection to the production well because the cold-water front is spread over a thicker reservoir interval. concluding remarks and perspectives the various geological and geophysical analyses presented here indicate that the gassum formation is suitable for geothermal exploitation in both of the prospect areas. the production may be further enhanced if geothermal energy is produced simultaneously from the gassum formation and from sandstones in the lower part of the overlying fjerritslev formation. this formation largely consists of tight mudstones, but in eastern sjælland its basal part contains several sandstones which may contribute to a geothermal production. although the simulations point out the eastern area as being more favourable for a geothermal production, other factors must also be considered in a final selection of a borehole location, such as drilling costs related to different drilling depths and non-geological parameters such as the position of the well in relation to the district heating and other surface infrastructure. well data from eastern sjælland are scarce and of varying quality. especially the lack of cores from penetrated sandstones is a shortcoming, as such material is very valuable in estimating reservoir properties as shown by the extrapolation of core data from the stenlille wells for the prospect areas in copenhagen. several packages of sandstones and intervening mudstones in the gassum formation have thus been correlated between the wells at stenlille and eastern sjælland and are therefore also expected to be present in the two prospect areas themselves. regional seismic mapping, palynological analysis and comparison of petrophysical log data patterns furthermore indicate that the paleogeographic setting and depositional environments during the deposition of the gassum formation were broadly similar in stenlille and in the prospect areas. in addition, u-pb radiometric dating of detrital zircon grains from the gassum formation indicates that all of these deposits were sourced mainly from reworking of the lower triassic bunter sandstone formation on the ringkøbing–fyn high, a regional basement ridge forming the southern margin of the danish basin. extrapolation of the stenlille data as far as to eastern sjælland inevitably implies some uncertainties. a new well in copenhagen from which cores, petrophysical log data and hydraulic test data can be collected and analysed will considerably increase the accuracy of predictions of reservoir properties of the gassum formation in greater copenhagen as well as in the hillerød and farum areas in northeastern sjælland, where initial investigations have also been performed. hence, the geological and economic risks associated with the establishment of a geothermal plant will be reduced, not only in copenhagen but in eastern sjælland as a whole. furthermore, a new well will make it possible to compare existing core data (including direct porosity and permeability measurements) with petrophysical log data and hydraulic test data from intervals of penetrated reservoir sandstone, and will thus provide a unique possibility to verify to what extent traditional petrophysical log data can be used to estimate the reservoir properties of geothermal sandstones. this knowledge is important for evaluation of the geothermal potential in a specific area based on data from existing wells, and for selecting suitable log tools for estimation of the porosity, permeability and injectivity of potential reservoir sandstones in general. acknowledgements the eudp is thanked for financial support. references balling, n., fuchs, s., poulsen, s.e., bording, t.s., nielsen, s.b., mathiesen, a. & nielsen, l.h. 2016: development of a numerical 3d geothermal model for denmark. proceedings, european geothermal congress, strasbourg, france. kristensen, l., hjuler, m.l., frykman, p., olivarius, m., weibel, r., nielsen, l.h. & mathiesen, a. 2016: pre-drilling assessments of average porosity and permeability in the geothermal reservoirs of the danish area. geothermal energy 4(6), 2–27, http://dx.doi.org/10.1186/ s40517-016-0048-6 vosgerau, h., mathiesen, a., andersen, m.s., boldreel, l.o., hjuler, m.l., kamla, e., kristensen, l., pedersen, c.b., pjetursson, b. & nielsen, l.h. 2016: a webgis portal for exploration of deep geothermal energy based on geological and geophysical data. geological survey of denmark and greenland bulletin 35, 23–26. authors’ address: geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: hv@geus.dk geological survey of denmark and greenland bulletin 33, 2015, 29-32 440 40 25 20 200 precipitation in filt rat ion 740 su rfa ce f lo w sea stream flow atmosphere evapotranspiration over sea unsaturated zone 15 abstraction 15 evapotranspiration groundwater wind groundwater flow 29 th e rapidly increasing impacts of climate change are likely to require changes in relevant institutions (ipcc 2012). an example is the growing need for immediate information on the entire water cycle (fig. 1), with quantitative assessments of critical hydrological variables and fl ow interactions between diff erent domains, e.g. atmosphere, plant-soil, surface water, groundwater and the sea, as they take place. potential measures include early warning systems, risk communication between decision makers and local citizens, sustainable land management, including land use planning as well as ecosystem management and restoration (ipcc 2012). early warning systems (kundzewicz 2013) that provide information and monitoring of past and present hydrological conditions as well as forecasting of hydrological conditions (e.g. water content, fl ow and water levels) are fi rst steps in developing fl ooding indicators for operational use. fig. 1. water balance in mm/year for a typical danish area. with climate change the freshwater cycle is no longer in a steady state. early warning and monitoring keeps us continuously updated and gives us an overview. this is important for our ability to combat the impact of climate change and manage the water resources proactively. it is possible that a nationwide system can be linked to local early warning systems and can make use of national overviews based on national hydrological models. flooding and drought events are complex phenomena with several key mechanisms including intense and long lasting precipitation or lack of precipitation in the case of drought and interaction with water uses. water consumers use groundwater for drinking water, food production, households and livestock, energy production and recreational activities. at the same time, water authorities have to manage surface and groundwater to sustain ecological systems, and support vital ecological conditions for plants and animals in rivers and wetlands. it is obvious that it is vital for society to have water in the right amount in the right place at the right time. th erefore, we must understand the water cycle, i.e. how overland drainage and surface runoff are generated, how water fl ows in the upper soil layers and in the deeper subsurface, how it is discharged in water courses and lakes and how freshwater interacts with the sea. one or two decades ago, water management was developed and operated under the assumption of static conditions, a paradigm which now is dead according to milly et al. (2008). it is stated by milly et al. (2008, p. 573) that “stationarity – the idea that natural systems fl uctuate within an unchanging envelope of variability – is a foundational concept that permeates a hydrological early warning system for denmark based on the national model hans jørgen henriksen, simon stisen, xin he and marianne b. wiese © 2015 geus. geological survey of denmark and greenland bulletin 33, 29–32. open access: www.geus.dk/publications/bull 3030 training and practice in water-resource engineering . . . . stationarity is dead because substantial anthropogenic change of the earth’s climate is altering the means and extremes of precipitation, evapotranspiration, and rates of discharge of rivers”. in the new paradigm, in the non-stationary world, real-time modelling and a continuity of monitoring systems are critical for dealing with the increasing impact of extremes. globally natural disaster costs have more than quadrupled since 1985 (georgieva 2014). early warning and monitoring systems that can transfer operational hydrological knowledge to community-based climate change adaptation planning and emergency management, are assumed to be fundamental for building societal resilience, both in the phases of pre-disaster, disaster response and post-disaster, and in general for extending monitoring techniques with network-based public participation. th e geological survey of denmark and greenland (geus) fi nanced a two year project (2012–2014), in order to determine the requirements for a hydrological system capable of providing real-time and forecast simulations based on a national hydrological model (the dk-model; www. vandmodel.dk; henriksen et al. 2003; stisen et al. 2012; højberg et al. 2013), and how the system can be linked with local early warning systems. th is paper describes four scenarios discussed at a participatory workshop in october 2014 and the workshop’s outcome. th e workshop provided geus with valuable insight and feedback relevant for the future development of a nation-wide, real-time modelling and water cycle monitoring system for denmark, including the possible input to an early warning system and real-time forecasting to operate at local level. structuring the needs – four scenarios for an early warning system design th e four scenarios, used in the workshop, for the design of a hydrological, real-time forecasting system linked to local level, community-based, emergency management are shown in fig. 2. scenario 1: an updated national model can provide an estimate of the actual status of water resources in denmark based on the calibrated national model (the dk-model) with updated climatic data, and available on-line with absolute values, indices or anomalies (below, same as, or above normal). scenario 2: a national model that can forecast the state of the water resources on a short timescale or as a seasonal prognosis, based on forward modelling simulations using weather forecast data. th is could be used as an early warning system and as a starting point for local, community-based emergency management. scenario 3: decision support is a scenario where local communities use a decision support tool for integrated assessment and management (kelly et al. 2013), for incorporating local knowledge and experience, and where simulation results from the national model in forecasting mode are included. th is scenario builds on a combination of monitored data and model output. th e output from this model might be with physical variables, thematic maps, indices or anomalies. scenario 4: a complex local model where a local detailed model (such as mike 11 or mike urban) uses simulated results from the national model in forecasting mode as input or boundary conditions. th is can be combined with monitoring data in a data assimilation framework. th e national model data in this scenario are physical variables delivered as pointor gridded data. a representative catchment area (skjern å; fig. 3) was selected for the workshop to test and demonstrate a prototype of a web-based, hydrological warning system for professional users. th e system provides hydrological simulations from groundwater levels, stream fl ow and water content in the root zone. webpages can be tailor-made to meet the requirements of end-users and continuously adapt to changing user demands. numerical results from simulations on a national fig. 2. the four real-time forecasting scenarios discussed in the text. scenario 1: updated dkmodel, scenario 2: dk-model forecast, scenario 3: local decision support and scenario 4: complex local model. 31 scale could also be provided. hourly values of climatic and discharge data at ahlergaarde station 250082 were used for the prototype demonstration. setting up the workshop a participatory workshop (hare & krykow 2005) was held in copenhagen for the purpose of eliciting stakeholder ideas and opinions and get feedback from prospective users. participants were planners and emergency managers from local authorities, water supply companies, ministries and consulting agencies. th e workshop had invited speakers from the holstebro, aarhus and fredensborg municipalities and local authorities that had already implemented local, real-time forecasting and early warning systems. th e danish hydraulic institute (dhi) and the danish meteorological institute (dmi), geus’ two partners in the project, also presented their experiences. both of these institutions have more than 20 years of experience developing and implementing modelbased early warning systems in denmark and abroad. prior to the workshop, an invitation was sent to those employed with climate change adaptation and emergency management at local and national levels. a questionnaire was attached to the invitation which explained why geus had initiated the project, introduced the four scenarios and briefly described the hydrological components that the national model can simulate. th e prototype of the web interface was described. th e goal of the workshop was an in-depth discussion of user requirements as a supplement to the web questionnaire which had 27 respondents. a total of 34 participants signed up for the workshop; eight came from geus, one from dmi, one from dhi, nine from local authorities, four from national ministries, ten from consulting companies and one from a large water-supply company. th e workshop was a one day event. th e fi rst block contained three presentations by geus participants: introduction to the real-time project, technical challenges in real-time modelling and web presentation of real-time data for the river skjern å catchment area. th is was followed by two invited presentations by dmi (better prediction of heavy rain) and the dhi (early warning systems in relation to hydrology and the freshwater cycle). th ree invited speakers from local authorities talked about fl ooding from the river storå in holstebro, the usserød å project in fredensborg and an early warning model with emphasis on fl ooding and groundwater monitoring in the urban area of aarhus. finally, there was a group session with three groups each addressing the same four questions. (1) what are the requirements for realtime forecasting? do you miss an overview of the hydrological state of an area in your daily work or in situations with fl ooding or drought? (2) what information would be benefi cial during such a situation, absolute or relative values; and which components of the hydrological cycle are the most important? (3) what is the time perspective in such an extreme event: days, hours or other? what would be the most appropriate frequency for updating the forecasts? (4) would a national overview make any diff erence in forecasting hydrological events, and how can or should the prototype be developed further? how can the present dk-model be part of such a system? outcome of the workshop th e results of the discussions were presented by groups. group 1 recommended linking national and local systems and presented a variety of requirements from diff erent local authorities posing potential challenges to the entire concept. th ey noted that problems with fl ooding, drought and emergency management are highly site specifi c. data assimilation is an important component in the early warning systems and there are many diff erent sources of observational data, e.g. from local authorities and the danish road directorate. a merged dataset is preferred. group 2 noted that real-time data for shallow groundwater levels are lacking but it may be possible to use observations from geotechnical boreholes. an investigation of geotechnical boreholes is necessary in order to establish a new operational groundwater level monitoring network. more emphasis should be given to a real-time early warning system instead of the present forecasting system with a short time horizon of only days. th e quality of early warnings should be quantifi ed in order to demonstrate how reliable the model can describe 5 km9°e 56°10´n discharge station hydraulic head observation stream ahlergaarde catchment station 250082 fig. 3. pilot case study for river skjern å, ahlergaarde catchment area. green: low elevation areas. dark yellow: high elevation areas. 3232 extremes. it is important to simulate water levels including storm surges. data and forecasts should be available online. group 3 noted that in many cases local authorities have their own local data which could provide input to the national model. local authorities may not have a groundwater model so co-operation between geus and local authorities are encouraged. local authorities experience an increased demand for warning and action capabilities. at the same time, local authorities are reluctant to issue warnings because they do not wish to be overcautious or risk subsequent claims. absolute values of physical variables are requested and data of levels of surface water and shallow groundwater are the most urgent. it was also suggested that continuous monitoring of water levels and early warning of changes in these levels are the most interesting for short-term forecasts, especially when local authorities do not have early warning systems, which is consistent with what was proposed by group 2. from a temporal perspective, precipitation events are highly diverse and it may not always be the short-term cloudbursts that are the most infl uential. long periods of rainfall (e.g. from a few days in succession to prolonged periods of rain) can signifi cantly increase groundwater levels in western and central jylland and snowmelting events can also cause extreme fl ooding. for the moment the national model has a limited applicability with its focus on water fl ow simulations. it is necessary to simulate water levels with local models. discussion and conclusion a participatory workshop discussion real-time forecasting was held to get feedback and get into dialogue with water planners and emergency managers from local authorities, water companies and national authorities. at the workshop, a prototype website illustrating four scenarios of national and local early warning systems was presented for the skjern å, ahlergaarde catchment area with selected events. th e workshop recommended that geus should focus on realtime modelling with the dk-model (scenario 1). th e fi rst step is to update the national coverage of climatic data input from dmi; real-time discharge fl ow and groundwater level monitoring are required; and the necessity of data assimilation and other types of uncertainty analyses have to be further evaluated. if a forecast model is included (scenario 2), complex data assimilation is required, however, this can compromise the water and mass balance of the model. th e water balance and simulation of the whole water cycle should be addressed by the early warning systems (most participants found that soil moisture, discharge in rivers and groundwater levels should be represented in such a system). eventually, an improved system for collecting observations of precipitation (or use of high-resolution radar measurements adjusted with rain-gauge data from a coarse network) is needed because the current network of rain-gauge stations has a poor national coverage. th e early warning system should be able to deliver results for discharge stations and boundary conditions for subsequent use in local models (scenarios 3 and 4). acknowledgements th e study was conducted as part of the geus-funded project: ‘realtidsvarsling’. th e paper is a nordress (nordress.hi.is) contribution. we are grateful for the input and feedback of the participants at the workshop. references georgieva, k. 2014: post-haiyan – a way forward. speech/14/441 by eu commissioner for international cooperation, humanitarian aid and crisis response. asem conference on disaster risk reduction and management. manila, 5 june 2014. hare, m. & krykow, j. 2005: participatory processes for the design of water storage areas – theme group iii inception report of the trust project. seecon report 1/2005, 68 pp. osnabrück: hoogheemraadschap van schielanden en der krimpenerwaard. henriksen, h.j., troldborg, l., nyegaard, p., sonnenborg, t.o., refsgaard, j.c. & madsen, b. 2003: methodology for construction, calibration and validation of a national hydrological model for denmark. journal of hydrology 280, 52–71. højberg, a.l., troldborg, l., stisen, s., christensen, b.s.b. & henriksen, h.j. 2012: stakeholder driven update and improvement of a national water resources model. environmental modelling and soft ware 40, 202–213. ipcc 2012: managing the risks of extreme events and disasters to advance climate change adaptation, a special report of working groups i and ii of the intergovernmental panel on climate change, 582 pp. cambridge: cambridge university press. kelly, r.a. et al. 2013: selecting among fi ve common modelling approaches for integrated environmental assessment and management. environmental modelling & soft ware 47, 159–181. kundzewicz, z.w. 2013: floods: lessons about early warning systems. in: gee, d. et al. (eds): late lessons from early warnings: science, precaution, innovation, 347–368. eea report no. 1/2013. copenhagen: european environment agency. milly, p.c.d., betancourt, j., falkenmark, m., hirsch, r.m., kundzewicz, z.w., lettenmaier, d.p. & stouff er, r.j. 2008: stationarity is dead: whither water management? science 319, 573–574. stisen, s., højbjerg, a.l., troldborg, l., refsgaard, j.c., christensen, b.s.b., olsen, m. & henriksen, h.j. 2012: on the importance of appropriate precipitation gauge catch correction for hydrological modelling at mid to high latitudes. hydrology and earth system sciences 16, 4157–4176. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: hjh@geus.dk geological survey of denmark and greenland bulletin 6, 9-18 9 jurassic syn-rift sedimentation on a seawards-tilted fault block, traill ø, north-east greenland henrik vosgerau, peter alsen, ian d. carr, jens therkelsen, lars stemmerik and finn surlyk middle–late jurassic rifting in east greenland was marked by westwards tilting of wide fault blocks bounded by major n–s-trending east-dipping synthetic faults. the syn-rift successions thicken westwards towards the faults and shallow marine sandstones show mainly southwards axial transport directions. an exception to this general pattern is found in south-east traill ø, which constitutes the e-tilted bjørnedal block, which is bounded to the west by the westwardsdipping antithetic vælddal fault. the stratigraphic development of the jurassic succession on this block shows important differences to the adjacent areas reflecting a different tectonic development. shallow marine sand seems initially to have filled accommodation space of the immediately adjacent block to the west. this block subsequently acted as a bypass area and much of the sediment was spilled eastwards onto the hangingwall of the east-dipping bjørnedal block. the succession on the bjørnedal block shows an eastwards proximal–distal decrease in sandstone– mudstone ratio, reflecting increasing water depth and progressive under-filling of the subbasin towards the east in agreement with the dip direction of the fault block. the transverse, mainly south-eastwards palaeocurrents, the eastwards increase in water depths and decrease in sandstone–mudstone ratio on the bjørnedal block are at variance with the standard picture of westtilted blocks with southwards-directed palaeocurrents and decrease in grain size. earlier palaeogeographic reconstructions have to be modified to account for the east-dipping hangingwall and different stratigraphic development of the area. the sea was thus open towards the east and there is no direct indication of a barrier or shoal east of traill ø. keywords: bjørnedal block, jurassic, north-east greenland, palaeocurrents, rifting, traill ø h.v.*, j.t.‡ & l.s., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. present addresses: *roskilde amt, køgevej 80, dk-4000 roskilde, denmark. ‡skude & jacobsen, næstvedvej 1, dk-4760 vordingborg. p.a. & f.s., geological institute, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: petera@geol.ku.dk i.d.c., oxford brookes university, headington, oxford ox3 0bp, uk. the main mesozoic rift phase in east greenland was initiated in mid-bajocian time, intensified during bathonian–oxfordian time, and culminated in the kimmeridgian–volgian. rifting was accommodated along major north–south-trending and east-dipping normal faults limiting wide westwards-tilted blocks. this resulted in the development of elongated fault-controlled marine embayments open to the south and with major rivers entering the northern heads of the embayments located in relay zones where the border faults shifted en échelon to the east (surlyk 1977a, 1978, 2003; surlyk & clemmensen 1983). transport of sand, silt and clay by marine currents was mainly axial towards the south. initial late bajocian progradation of shallow marine sands reached the southern end of the exposed basin. the sandy system (pelion formation) geological survey of denmark and greenland bulletin 5, 9–18 © geus, 2004 geus bulletin no 5.pmd 29-10-2004, 11:139 10 g r ee n la n d ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ geographical society ø traill ø jameson land 24ºw 23ºw 22ºw 72º45′n 72º30′n 72º15′n 72º00′n n n svinhufvud bjerge rold bjerge 10 km mols bjerge kong oscar fjord månedal mestersvig palaeogene intrusives palaeogene extrusives cretaceous jurassic triassic upper permian carboniferous devonian pre-devonian fault vælddal bjørnedal drømmebugten section b section a bf mf mbf vf bf bordbjerg fault mf månedal fault mbf mols bjerge fault vf vælddal fault sedimentological sections steenstrup dal forårsdal ■■ 72º15'n 23ºw 23ºw 72º15'n 54 1–5 3 1 2 bj ør ne da l st ee ns tr up d al d rø m m eb ug te n b a b g re en la nd fo rå rs da l 73º00′n geus bulletin no 5.pmd 29-10-2004, 11:1310 11 progressively backstepped during the bathonian and was eventually drowned in the late callovian due to increased rates of extension and a long-term eustatic rise in sea level (engkilde & surlyk 2003; surlyk 2003). an exception to this simple pattern is found in southeastern traill ø, where an east-dipping fault block, more than 30 km wide, was formed during early rifting (figs 1, 2; carr 1998). it was limited to the west by the west-dipping vælddal fault (donovan 1953; carr 1998). the block, which is here termed the bjørnedal block, probably continues southwards into northern jameson land where poorly exposed east-dipping correlative strata occur on the south side of kong oscar fjord (fig. 1). the aim of this study is to compare and contrast the syn-rift middle–late jurassic stratigraphic syn-rift development on the eastwards-dipping bjørnedal block on south-east traill ø with the rest of the jurassic basin of east greenland, which is characterised by westwards-tilted blocks. nw se nw se 1000 syenitic plutons 500 0 m 1000 m 5 km 500 0 bo rd bj er g fa ul t m ån ed al f au lt st ee ns tr up d al v æ ld da l v æ ld da l f au lt m ån ed al fa ul t zo ne fo rc hh am m er d al d rø m m eb ug te n fo rå rs da l m ol s bj er ge f au lt basaltic sills cretaceous upper jurassic middle jurassic triassic upper permian carboniferous fault, arrows indicate direction of movement rold bjerge – mols bjerge (section b) south coast, traill ø (section a) ? bjørnedal block ? facing page: fig. 1. map of traill ø area showing the main faults and outcrops of the jurassic succession. positions of the structural cross sections (a, b) on fig. 2 are indicated. inset map (b) shows location of measured sedimentological sections (1–5). fig. 2. nw–se-oriented cross-sections showing the geological structure of eastern traill ø. note the west-dipping vælddal fault, which forms the western limit of the bjørnedal block. virtually all other mesozoic faults in east greenland dip towards the east and delimit westwards-tilted blocks. the positions of the sections are indicated on fig. 1. based on koch & haller (1971). geus bulletin no 5.pmd 29-10-2004, 11:1411 12 the south coast of traill ø upper palaeozoic – triassic deposits are exposed along the western part of the south coast of traill ø, whereas jurassic and younger sediments are restricted to the eastern part (fig. 1). the jurassic succession is bounded to the west by the månedal fault and is cut by the vælddal fault situated 15 km further to the east (figs 1, 2). the peninsula east of the vælddal fault is made up mainly by the palaeogene kap simpson syenite complex which extends for about 30 km in a nw–se direction parallel to the coast (fig. 1). the coastal cliffs are high and steep, and numerous palaeogene sills and dykes intrude the jurassic succession. it has accordingly received very little attention. it was assigned to the yellow and black series by donovan (1953) and this was followed in the geological map of koch & haller (1971). the yellow series was very loosely defined and covered middle and lower upper jurassic sandstone-dominated successions. on traill ø, it includes the pelion and olympen formations of current usage. the black series includes upper jurassic dark grey mudstones and black shales corresponding in part to the bernbjerg formation (surlyk 1977b). successively younger jurassic strata are exposed from west to east along the coast east of the vælddal fault (fig. 2). the succession includes lower bajocian sandstones and mudstones of the bristol elv formation, upper bajocian – lower callovian sandstones of the pelion formation, callovian mudstones of the fossilbjerget formation, intercalated with sandstones of the parnas member (top member of the pelion formation), overlain by lower–middle oxfordian mudstones and sandstones of the olympen formation, and upper oxfordian – kimmeridgian dark grey and black mudstones of the bernbjerg formation (fig. 3). the exposures along the coastal cliffs vary in quality. from a stratigraphic point of view they are generally good, whereas sedimentary structures commonly are obliterated by the effects of the palaeogene sills and dykes. for this study, five sedimentological sections were measured along the coast allowing construction of a w–e dip section (fig. 3). the sections are correlated by lateral tracing of major sedimentary packages and drowning surfaces in the field and on a high-quality photo-mosaic, and by ammonite dating of a number of levels. the succession dips about 13° to the east except in the easternmost section where it dips 9° towards the north, but in this area dip-directions reflect disturbances by the immediately adjacent syenite complex (fig. 1). the oldest strata are exposed to the west immediately east of the vælddal fault where jurassic strata rest directly on triassic redbeds of the fleming fjord formation (stauber 1942). the base of the middle jurassic sandstones is not exposed further east, but fluvial to marginal marine sandstones and mudstones of the bristol elv formation (price & whitham 1997; stemmerik et al. 1997; therkelsen & surlyk 2004, this volume) occur below marine upper bajocian sandstones of the pelion formation in two sections. three of the measured sections (fig. 3, sections 1–3) reach up into black mudstones of the upper oxfordian – kimmeridgian bernbjerg formation. stratigraphic development the basal part of the jurassic succession consists of channelised, trough cross-bedded pebbly sandstones interbedded with dark grey laminated shales occasionally with thin sandstone ripples. crude fining-upwards trends can be recognised, and rootlet beds are common. palaeocurrents are mainly towards the south-west (fig. 3). the interbedded shales contain abundant leaves and scattered tree trunks. both facies contain scattered trace fossils of marine affinity. the sandstones are interpreted as having been deposited in coastal rivers and the shales were formed by drowning of the fluvial system either due to delta switching and abandonment or to base-level rise. the trace fossils suggest some marine influence and deposition of the shales probably took place under estuarine conditions or in interdistributary bays. a succession of fluvial deposits at the base of the jurassic succession of traill ø was recognised independently by facing page: fig. 3. nw–se-oriented dip-parallel correlation panel of the bjørnedal block, based on five measured sections along the south coast of traill ø (for location, see fig. 1b). note the eastwards, down-dip decrease in sandstone–mudstone ratio and increase in the thickness of mudstone packages. the key motif is a coarsening-upwards parasequence composed of offshore mudstones, overlain by offshore transition zone heteroliths followed by shoreface sandstones. the parasequences stack into parasequence sets, numbered ps1–8 from below. the succession belongs from below to the dominantly fluvial bristol elv formation, the marine sandstone-dominated pelion formation, the mudstone-dominated fossilbjerget formation, the olympen formation of interbedded marine sandstones and mudstones, and the dark, deeper marine mudstone-dominated bernbjerg formation. p., pelion. geus bulletin no 5.pmd 29-10-2004, 11:1412 13 pe lio n fm br is to l e lv f m o ly m pe n fm be rn bj er g fm vælddal (section 2) ? e. steenstrup dal (section 3) bjørnedal (section 1) e. steenstrup dal (section 4) ? mud sand f m c forårsdal (section 5) 100 m mud sand f m c mud sand f m c mud sand f m c mud sand f m c mud sand f m c ps8 ps7 ps6 ps5 ps4 ps3 ps2 ps1 8 km 10 km 2 km 19 km mudstone lithology siltstone sandstone pebbles concretions massive structures wavy bedding plane bedding plane lamination hummocky crossstratification trough cross-bedding single trough cross-beds planar cross-bedding diplocraterion habichi trace fossils curvolithos multiplex ophiomorpha nodosa degree of bioturbation fossils and symbols belemnite ammonite bivalve roots palaeocurrent direction coalified wood parasequence setps1–8 lithostratigraphic boundary parasequence set boundary plant fragments c. pompeckji c. apertum c. nordenskjoeldi c. apertum c. apertum a. serratum a. cymodoce c. pompeckji c. pompeckji a. ishmae a. cranocephaloide parnas mb (p. fm) fossilbj. fossilbj. fm g e u s b ulletin no 5.pm d 29-10-2004, 11:14 13 14 price & whitham (1997) and stemmerik et al. (1997) and is placed in the new bristol elv formation by therkelsen & surlyk (2004, this volume), who refer it tentatively to the early bajocian. the deposits described here from the south coast of traill ø are also referred to the bristol elv formation on the basis of the dominant pebbly sandstone lithology and the fluvial style of deposition. the boundary between the bristol elv formation and the overlying marine sandstones of the pelion formation is not exposed. the bristol elv formation is overlain by marine sandstones and thin mudstones of the pelion, fossilbjerget and olympen formations, which are composed of coarsening-upwards units a few to several tens of metres thick. these units start with dark grey laminated to bioturbated offshore mudstones with thin subordinate beds of silty fine-grained sandstone or mediumgrained sandstone, overlain by heterolithic deposits, which give way to mediumto coarse-grained trough and planar cross-bedded sandstones. the boundary between the mudstones and the sandstones is usually gradational but is sharp and erosional in a few cases. the transitional interval shows hummocky cross-stratification in some sections. the top of the coarseningupwards units is a sharp drowning surface, overlain by offshore mudstones of the next unit, occasionally with an intervening pebble lag composed of subrounded to rounded quartzite pebbles up to 3 cm long. metre-long u-burrows of diplocraterion habichi extend downwards from the drowning surface. the coarsening-upwards units were formed by progradation of shoreface sands across offshore transition zone heteroliths and offshore mudstones and are bounded by drowning surfaces. in some cases these surfaces have been modified by transgressive shoreface erosion and may conceal subtle sequence boundaries. the thickest and coarsest pebble lag, which occurs in the most distal section, suggests bypass of the beach during sea-level fall and subsequent transgressive shoreface winnowing and erosion. the units thus represent parasequences or simple sequences, but in this context they are for simplicity termed parasequences throughout. in some of the parasequences, the sandstone part has a sharp base formed by shoreface erosion during progradation under sea-level fall and interpreted as a forced regressive surface of erosion. the parasequences stack in parasequence sets, which are numbered ps1–8 (fig. 3). a total of six parasequence sets are recognised in the pelion formation. parasequence set 6 is much finer olympen f/p pelion ps3 ps4 ps5 ps6 ps7 ps8 section 3 section 4 section 4 100 m syenite wnw ese bernbjerg cretaceous fig. 4. the coastal section of the bjørnedal block, immediately east of steenstrup dal, showing yellow sandstones of the pelion formation below, overlain at a pronounced drowning surface by dark grey mudstones of the fossilbjerget formation followed by alternating sandstones and mudstones of the olympen formation, and finally by dark grey mudstones of the bernbjerg formation. the height of the cliff is approximately 1000 m. ps3–8, parasequence sets; stippled lines, major flooding surfaces; f/p, fossilbjerget and pelion (parnas mb) formations. geus bulletin no 5.pmd 29-10-2004, 11:1414 15 grained than the lower parasequence sets and lithostratigraphically it represents interfingering between mudstones of the fossilbjerget formation and the lower callovian parnas member of the pelion formation (alsen & surlyk 2004, this volume). the pelion–fossilbjerget parasequence sets thus constitute a composite aggradational to retrogradational package (figs 3, 4). sandstones and mudstones of the olympen formation overlie the interfingering pelion–fossilbjerget couplet. this unit has not yielded any ammonites in the studied sections, but is elsewhere of early–middle oxfordian age (surlyk 1978; price & whitham 1997) and this age assignment is corroborated by the bracketing callovian age of the top of the fossilbjerget formation and late oxfordian base of the overlying bernbjerg formation. the olympen formation is composed of about five stacked coarsening-upwards units, which are thinner and finer grained than those of the pelion formation. they are interpreted as parasequences and form two parasequence sets, ps7–8, the lower of which includes the uppermost mudstones of the fossilbjerget formation (fig. 3). this development is similar to the type area in central jameson land, where the formation also comprises two thick sandstone units and an intervening mudstone unit (larsen & surlyk 2003). the olympen formation is overlain by black offshore mudstones with a few thin sandstones of the upper oxfordian – kimmeridgian bernbjerg formation, reflecting final drowning of the sandy depositional systems. the marine middle and upper jurassic succession of south-east traill ø thus shows a stepwise backstepping trend. it commences with the lower aggradational to retrogradational parasequence sets of the pelion formation dominated by rather coarse-grained sandstones topped by the intertonguing fossilbjerget formation mudstones and finer grained sandstones of the parnas member (pelion formation). then follows the overall finer grained olympen formation, which is overlain by the bernbjerg formation mudstones. the main backstepping events can be roughly dated to the callovian–oxfordian and middle–late oxfordian boundaries. major drowning surfaces the drowning surfaces separating the thicker parasequence sets can be traced between the sections. it is difficult to evaluate their significance and regional extent but the drowning surfaces topping parasequence sets 3 and 5 are well dated palaeontologically and seem to represent regional flooding events (figs 3, 4). cranocephalites pompeckji thus characterises one such event, which can be traced from southern jameson land to traill ø and possibly as far north as hold with hope (vosgerau et al. 2004, this volume). the c. pompeckji chronozone is upper bajocian, possibly reaching up into the lowermost bathonian (callomon 1993). cadoceras apertum, which marks the base of the callovian (i.e. = c. apertum chronozone of callomon 1993, p. 103) is found 200 m above c. pompeckji. it seems likewise to represent a major regional flooding event (piasecki & larsen 1998; engkilde & surlyk 2003). palaeocurrents fluvial and marine palaeocurrents in the middle jurassic deposits of east greenland are mainly axial towards the south with a subordinate northwards tidal component. the bjørnedal block represents a significant exception to this pattern. the fluvial bristol elv formation shows palaeocurrents towards the south-west in the western part of the dip transect (figs 3, 5). the palaeocurrents of the overlying marine sandstones of the pelion and olympen formations are, however, mainly towards the south-east and north-east (figs 3, 5), and marine currents thus essentially moved down the hangingwall slope in an offshore direction. a subordinate sw–ne-oriented tidal system is recorded in the easternmost section (figs 3, 5). down-dip facies development on the bjørnedal block a marked change in facies is recorded down the hangingwall of the bjørnedal block (fig. 3). the up-dip western sections are more sand-rich, and mudstone units are relatively thin. down-dip, the mudstone units become thicker and the whole succession expands in thickness. the interval from the upper bajocian cadoceras pompeckji chronozone to the lower callovian c. apertum chronozone is thus about 200 m thick towards the west and increases to at least 250 m at the eastern down-dip end of the section. the thicknesses are measured between the correlative drowning surfaces, and the down-dip thickness increase thus amounts to 50 m over 20 km. the western up-dip area is thus slightly condensed compared to the down-dip area, and was bypassed by much of the finer-grained geus bulletin no 5.pmd 29-10-2004, 11:1415 16 sediment, which was deposited on the lower parts of the hangingwall slope where creation of accommodation space was greater. comparison of stratigraphy on wand e-dipping blocks the eastwards-dipping bjørnedal block contains thick mudstone units at the base of the parasequence sets. this contrasts markedly with the pelion formation parasequences on the wide westwards-dipping blocks elsewhere in east greenland, which consist almost exclusively of sandstones (engkilde & surlyk 2003). the marine palaeocurrents are mainly towards the south-east on the bjørnedal block, whereas they are towards the south or ssw on the westwards-dipping blocks. the succession similarly shales out towards the south-east on the bjørnedal block and towards the s–ssw on the latter blocks. the most important difference is that sediment influx to the bjørnedal block was derived from bypass and overspill of the adjacent block to the west, as also noted by carr (1998), whereas sediment influx to the west-dipping blocks was directly from rivers at the heads of the structurally controlled embayments. stratigraphic implications the marine middle jurassic deposits in adjacent parts of east greenland, notably in jameson land are placed in the proximal sandy pelion formation and the distal mudstone-dominated fossilbjerget formation. there is little interfingering between the two formations except for the upper part where the uppermost tongue of the pelion formation (parnas member) is intercalated within the top fossilbjerget formation (heinberg & birkelund 1984; engkilde & surlyk 2003). the stratigraphic development of the bjørnedal block differs in its regular alternation between mudstones and sandstones, which form a composite aggradational to retrogradational stack of parasequence sets culminating in the black mudstones of the bernbjerg formation. a pragmatic solution to this stratigraphic problem is to place the lower sandstone-dominated yellow sandstone package of parasequence sets 1–5 in the pelion formation. the overlying dark grey mudstone (lower part of parasequence set 6) is placed in the fossilbjerget formation and the overlying sandstone of parasequence set 6 in the parnas member of n = 39 circle = 15% n = 40 circle = 15% vælddal and e. steenstrup dal forårsdal vælddal and e. steenstrup dal forårsdal vælddal, e. steenstrup dal and forårsdal n = 12 circle = 15% pelion fm bristol elv fm olympen fm n = 10 circle = 30%n = 5 circle = 60% fig. 5. palaeocurrent roses for the bristol elv, pelion and olympen formations. measurements of cross-bed foreset dip azimuths. note the dominance of e–se down-dip directions which contrasts markedly with the dominant s–ssw directions of the pelion and olympen formations elsewhere in east greenland. geus bulletin no 5.pmd 29-10-2004, 11:1416 17 the pelion formation. the mudstone–sandstone boundary in parasequence set 6 is characteristically sharp and is well suited as a lithostratigraphic boundary. the mudstone of the lower part of parasequence set 7 forms the top tongue of the fossilbjerget formation (alsen & surlyk 2004, this volume). the olympen formation comprises the sandstone-dominated upper part of parasequence set 7 and parasequence set 8 (fig. 3). the ages of the three formations, as here defined, fit well with other areas in east greenland. the base of the pelion formation is not exposed or has not yielded any fossils but is thought to belong to the upper bajocian cadoceras borealis chronozone, in agreement with evidence from nearby areas in central traill ø. parasequence sets 4 and 5 from the upper part of the formation belong to the c. pompeckji and c. apertum chronozones, and the top of the formation (parnas member) falls in the basal part of the c. nordenskjoeldi chronozone (fig. 3). the lower wedge of the fossilbjerget formation is only 25 m thick and belongs to the c. apertum chronozone. the age of the top wedge of the fossilbjerget formation (base of parasequence set 7) is poorly constrained but is probably still callovian. this development is similar to that found in nearby bjørnedal (alsen & surlyk 2004, this volume) and in central jameson land (engkilde & surlyk 2003). palaeogeographic implications the south-eastwards down-dip palaeocurrent directions and proximal to distal facies changes of the jurassic succession on the bjørnedal block indicate that the sea was open and deepest towards the east (fig. 6). this contrasts with the adjacent parts of the middle jurassic basin of east greenland which are characterised by n–s-oriented marine embayments limited to the east by elongated peninsulas, islands or submarine shoals formed by uplifted crests of westwardstilted blocks. these embayments were open and deepest towards the south as shown by overall southwardsdirected palaeocurrents and decrease in grain size (surlyk 1977b, 1978, 2003; surlyk & clemmensen 1983). conclusions the bjørnedal block was formed during rifting initiated in late bajocian time (carr 1998). the block is bounded to the west by the west-dipping vælddal fault and is tilted towards the east in contrast to the middle– 50 km 20°w24°w28°w 73°n 72°n 71°n traill ø geographical society ø jameson land sandstone and sandy mudstone sandstone sandstone with conglomerate land silty mudstone inferred coastline hypothetical coastline vælddal fault main direction of sediment transport fig. 6. generalised bathonian (middle jurassic) palaeogeography and facies distribution in the geographical society ø, traill ø and jameson land region. the sea was open towards east in the area comprising south-eastern traill ø and north-eastern jameson land due to the east-dipping nature of the more than 30 km wide bjørnedal block, which was formed during early rifting. this contrasts with the remaining parts of the jurassic basin of east greenland where n–s-oriented marine embayments were limited to the east by elongated peninsulas, islands or submarine shoals formed by uplifted crests of westwards-tilted blocks. based on surlyk (1977b). late jurassic west-dipping blocks characterising the rest of the jurassic basin of east greenland. the middle – lower upper jurassic succession in east greenland shows mainly axial, southwards-directed palaeocurrents and sediment entered the basins at the head of faultcontrolled embayments. the succession on the bjørnedal block differs in showing eastand southeast-directed palaeocurrents and associated proximal geus bulletin no 5.pmd 29-10-2004, 11:1417 18 to distal facies changes, transverse to the axis of the rift basin. the middle jurassic sediments bypassed the block west of the vælddal fault (carr 1998) and spilled over onto the eastwards-dipping hangingwall of the bjørnedal block. existing palaeogeographic reconstructions have thus been modified to account for the eastwards dip of the bjørnedal block. in this area the sea was open and deepened towards the east and there is no indication of a barrier or shoal to the east. acknowledgements this paper is a contribution to the project ‘resources of the sedimentary basins of north greenland and east greenland’ that is supported by the danish research councils. we are grateful to gregers dam and michael larsen for constructive reviews. references alsen, p. & surlyk, f. 2004: maximum middle jurassic transgression in east greenland: evidence from new ammonite finds, bjørnedal, traill ø. in: stemmerik, l. & stouge, s. (eds): the jurassic of north-east greenland. geological survey of denmark and greenland bulletin 5, 31–49 (this volume). callomon, j.h. 1993: the ammonite succession in middle jurassic of east greenland. bulletin of the geological society of denmark 40, 83–113. carr, i.d. 1998: facies analysis and reservoir characterisation of jurassic sandstones from bjørnedal, central east greenland, 245 pp. unpublished ph.d. thesis, institute for sedimentology, university of reading, uk. donovan, d.t. 1953: the jurassic and cretaceous stratigraphy and palaeontology of traill ø, east greenland. meddelelser om grønland 111(4), 150 pp. engkilde, m. & surlyk, f. 2003: shallow marine syn-rift sedimentation: middle jurassic pelion formation, jameson land, east greenland. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 813–863. heinberg, c. & birkelund, t. 1984: trace-fossil assemblages and basin evolution of the vardekløft formation (middle jurassic, central east greenland). journal of paleontology 58(2), 362–397. koch, l. & haller, j. 1971: geological map of east greenland 72º–76ºn. lat. (1:250 000). meddelelser om grønland 183, 26 pp., 13 maps. larsen, m. & surlyk, f. 2003: shelf-edge delta and slope deposition in the upper callovian – middle oxfordian olympen formation, east greenland. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 931–948. piasecki, s. & larsen, m. 1998: biofacies and sequence stratigraphy in an intracratonic seaway, middle to upper jurassic, east greenland. abstracts (cd-rom), american association of petroleum geologists 1998 annual meeting, salt lake city (utah), 2 pp. price, s.p. & whitham, a.g. 1997: exhumed hydrocarbon traps in east greenland: analogs for the lower–middle jurassic play of northwest europe. american association of petroleum geologists bulletin 81, 196–221. stauber, h. 1942: die triasablagerungen von ostgrönland. meddelelser om grønland 132(1), 325 pp. stemmerik, l., clausen, o.r., korstgård, j., larsen, m., piasecki, s., seidler, l., surlyk, f. & therkelsen, j. 1997: petroleum geological investigations in east greenland: project ‘resources of the sedimentary basins of north and east greenland’. geology of greenland survey bulletin 176, 29–38. surlyk, f. 1977a: mesozoic faulting in east greenland. in: frost, r.t.c. & dikkers, a.j. (eds): fault tectonics in n.w. europe. geologie en mijnbouw 56, 311–327. surlyk, f. 1977b: stratigraphy, tectonics and palaeogeography of the jurassic sediments of the areas north of kong oscars fjord, east greenland. bulletin grønlands geologiske undersøgelse 123, 56 pp. surlyk, f. 1978: jurassic basin evolution of east greenland. nature 274(5667), 130–133. surlyk, f. 2003: the jurassic of east greenland: a sedimentary record of thermal subsidence, onset and culmination of rifting. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 659–722. surlyk, f. & clemmensen, l.b. 1983: rift propagation and eustacy as controlling factors during jurassic inshore and shelf sedimentation in northern east greenland. sedimentary geology 34, 119–143. therkelsen, j. & surlyk, f. 2004: the fluviatile bristol elv formation, a new middle jurassic lithostratigraphic unit from traill ø, north-east greenland. in: stemmerik, l. & stouge, s. (eds): the jurassic of north-east greenland. geological survey of denmark and greenland bulletin 5, 19–29 (this volume). vosgerau, h., larsen, m., piasecki, s. & therkelsen, j. 2004: a new middle–upper jurassic succession on hold with hope, north-east greenland. in: stemmerik, l. & stouge, s. (eds): the jurassic of north-east greenland. geological survey of denmark and greenland bulletin 5, 51–71 (this volume). geus bulletin no 5.pmd 29-10-2004, 11:1418 e2019430208-01 the landslide of 17 june 2017 at karrat fjord, central west greenland, triggered a tsunami that caused four fatalities. the catastrophe highlighted the need for a better understanding of landslides in greenland and initiated a recent nation-wide landslide screening project led by the geological survey of denmark and greenland (geus; see also svennevig (2019) this volume). this paper describes an approach for compiling freely available data to improve geus’ capability to monitor active landslides in remote areas of the arctic in near real time. data include seismological records, spaceborne synthetic aperture radar (sar) data and multispectral optical satellite imagery. the workflow was developed in 2018 as part of a collaboration between geus and scientists from the technical university of denmark (dtu). this methodology provides a model through which geus will be able to monitor active landslides and provide relevant knowledge to the public and authorities in the event of future landslides that pose a risk to human life and infrastructure in greenland. we use a minor event on 26 march 2018, near the site of the karrat 2017 landslide, as a case study to demonstrate 1) the value of multidisciplinary approaches and 2) that the area around the landslide has continued to be periodically active since the main landslide in 2017. geological setting and description of the landslide area the geology of the karrat area (fig. 1) predominantly consists of reworked archaean gneiss interfolded with supracrustal rocks of the palaeoproterozoic karrat group (henderson & pulvertaft 1967). locally, around the landslide and surrounding unstable areas, the succession consists of archaean gneiss overlain by palaeoproterozoic semipelitic to pelitic schist. the geology of the karrat 2017 landslide has not been mapped in detail; however, an ongoing mapping project aims to update the geological maps of the area (e.g. sørensen & guarnieri 2018). during the landslide of 17 june 2017, 35–58 million m3 of material were mobilised (bessette-kirton et al. 2017; gauthier et al. 2018). a preliminary examination of the karrat 2017 landslide shows that a number of precursor events took place in the years prior to the main landslide. unstable areas that may fail in the future have also been identified. of these, two principal areas (marked with y and z in fig. 2a) west of the scarp of the main slide (marked by ‘x’ in fig. 2a) a multidisciplinary approach to landslide monitoring in the arctic: case study of the march 2018 ml 1.9 seismic event near the karrat 2017 landslide kristian svennevig*1, anne munck solgaard1, sara salehi1, trine dahl-jensen1, john peter merryman boncori2, tine b. larsen1 and peter h. voss1 geus bulletin is free to individuals and institutions in electronic form. the author(s) retain copyright over the article contents. research article | open access geus bulletin vol 43 | e2019430208 | published online: 01 july 2019 https://doi.org/10.34194/geusb-201943-02-08 fig. 1 overview map of the uummannaq fjord area showing the location of the 26 march 2018 1.9 ml seismic event (ellipse) along with the position of the 17 june 2017 landslide at karrat and the nearby instabilities. the size of the landslide and nearby instabilities are slightly exaggerated for clarity on the small map scale. seismic stations in uummannaq and nuugaatsiaq are shown along with settlements in the area. fig. 2 71°30’n 71°00’n 51°w u u m m a n n a q f j o r d siggnup nunaa n u u s s u a q nuugaatsiaq illorsuit ukkusissat marmorilik qaarsut niaqornat uummannaq saattut g r e e n l a n d i c e s h e e t 53°w fig. 1 k a r r a t f j o r d seismic station location of event ml 1.9 (26 march 2018) landslide of 17 june 2017 instability settlement g re en lan d 25 km https://doi.org/10.34194/geusb-201943-02-08 e2019430208-02 show signs of significant deformation both before and after the 2017 landslide. this activity is observable in both optical satellite images and in the results derived from remotely sensed differential synthetic aperture radar (sar) interferometry (dinsar; y and z in figs 2a, d, e). data the various data sources used here have different temporal and spatial resolutions. individually, they provide unique information for landslide monitoring, but their value increases significantly when combined. as such, integration of data from seismograms, dinsar and optical images can provide a more complete understanding of the geological processes contributing to landslide activity. seismic events the present greenland ice sheet monitoring network (glisn) of seismographs consists of 21 stations, situated about 200 km or more apart (clinton et al. 2014). given this large distance, the uncertainty associated with the horizontal location of detected earthquakes or other types of seismic events in greenland can be up to 50 km. differentiation of tectonic and non-tectonic events is based on the judgement of an experienced seismologist. for example, the main seismic features associated with tectonic earthquakes are clearly separated body waves, higher frequency content and a well-located hypocentre at depth, whereas non-tectonic events are not dominated by higher frequencies. these can be caused in several ways. glacial events, such as those caused by calving ice, have an epicentre located near an outlet glacier and often contain a low frequency component (ekström et al. 2003; nettles et al. 2008). large landslides can also generate seismic signals, but these are often of longer duration than those caused by tectonic earthquakes, such as the seismic signals caused by the 2000 paatuut (dahljensen et al. 2004) and the 2017 karrat (clinton et al. 2017) landslides. for smaller landslide events (e.g. events without catastrophic failure and rock avalanche activity), the duration will be shorter and the amplitude smaller, and it may be difficult to distinguish them with seismic data alone from other non-tectonic seismic sources, e.g. signals generated by moving sea ice or glaciers. integration of other data sources are thus necessary. optical satellite images sentinel-2a and -2b are multispectral optical satellite imaging systems that cover 13 spectral bands at various spatial resolutions: 4 bands at 10 m resolution (including visual light), 6 bands at 20 m and 3 bands at 60 m. the revisit frequency of sentinel-2a and -2b over the karrat fjord area is 1–2 days at around 15:30–16:00 utc (12:30–13:00 local time). however, no images are collected between the end of october and the start of march as it is too dark. the present study is limited to visual interpretation of sentinel-2 images as change detection algorithms have not been implemented. dinsar dinsar (rosen et al. 2000) was applied to sentinel-1a and sentinel-1b synthetic aperture radar (sar) data acquired between 24 february and 13 april 2018. this method provides one-dimensional ground motion measurements in the satellite line-of-sight direction, i.e. towards and away from the radar. the main acquisition mode of sentinel-1 over land is the interferometric wide (iw) swath mode, which provides 250 km × 250 km images at a 5 m × 20 m spatial resolution in the ground range and flight-path directions, respectively. the karrat area is covered by two satellite tracks: descending track 25 and ascending track 90. however, on ascending passes in track 90, the area of interest slopes towards the radar, leading to significant radar foreshortening (and in some locations even to layover). in turn, this causes geometric decorrelation of the radar signal even for small interferometric baselines, due to the horizontal spreading of the scatterers within each resolution cell (i.e. due to the poor line-of-sight resolution). for this reason, we used data from track 25 only in this analysis. unlike optical data from e.g. sentinel-2, sar has the advantage of being insensitive to cloud cover and solar illumination, and imagery can therefore be acquired year-round. dinsar can be applied to data collected along the same radar track, which are available every six days for the sentinel1a/b constellation on the greenland ice sheet margin. one of the main requirements of dinsar is a sufficient level of coherence (i.e. statistical similarity) between the two acquisitions. in practice, this limits the number of good interferograms collected during the winter season, as snowfall may change the surface morphology between two acquisitions and cause loss of coherence. for this study, both 6 and 12 day differential interferograms were constructed. topographic contributions to the interferometric phase were removed using arcticdem version 2.0 (porter et al. 2018) and additional corrections (courtesy of e.v. sørensen, geus) were e2019430208-03 plot start time: 2018 26 march 21:20 56 21:21 seconds time (utc) 100 20 30 40 50 nuug hhz nuug hh1 nuug hh2 a b c fig. 2 x y z 2018 26 march 15:29 41 1 km2018 27 march 15:48 59 250 m x y z x y z d e march 20 to 26 2018 march 26 to april 1 2018 fig. 2 composite figure with the different datasets used to constrain the event. a: sentinel-2a rgb image from 26 march 2018 15:29 41 utc. dashed line at x shows the scar of the 17 june 2017 landslide and y shows the outline of the active area to the sw. z indicates the position of a large older landslide with periodic activity in the toe. b: sentinel-2b rgb image from 27 march 2018 15:48 59 utc. red arrows show rock fall at the unstable area y and the green arrow shows rock fall at the back scarp of the karrat 2017 event. the dotted black line shows the outline of the area affected by the event, as observed in the inset image. c: the 26 march 2018 21:21 utc 1.9 ml seismic event as recorded on the nuugaatsiaq seismograph (fig. 1). d: sentinel-1 differential interferogram from before the event, (20180320-20180326), the latter interferogram was acquired only hours prior to the event. x, y and z refer to the same areas in a. e: sentinel-1 differential interferogram spanning the event from (20180326-20180401). x, y and z refer to the same areas in a. e2019430208-04 made to account for the height variations associated with the june 2017 landslide. the 26 march 2018 ml 1.9 event on 26 march 2018 at 21:21 utc (18:21 local west greenland time) a shallow, low frequency non-tectonic ml 1.9 seismic event was recorded on several seismographs in west greenland (figs 1, 2c; see the earlier discussion of tectonic vs. non-tectonic event features). the event was located to an ellipsoid of c. 100 by 20 km covering an area of known landslide activity near the 2017 karrat landslide. the time between the p and s wave arrivals corresponded to the distance between the seismic station in nuugaatsiaq, and the karrat landslide area. the ellipsoid also encompassed two glacial outlets (fig. 1). to investigate the source of the seismic signal, we acquired sentinel-2a and 2b images from 26 march 15:29 utc and 27 march 15:48 utc. comparison of the two images show dark colouration of the snow (fig. 2b) indicating rock fall from the known active area (red arrows in fig. 2b) and the whole length of the headscarp of the karrat 2017 landslide (green arrow in fig 2b). the rock fall was only observed very locally on the coast and a screening of neighbouring slopes and glacial outlets showed no changes between the two images. temperatures were well below freezing during the whole period. images from the week before and after the event were also screened and showed no increased rock fall activity. in the period leading up to the event, the 6 day pairs of sar images are coherent in the whole slope area, and the differential interferograms show activity constrained to the 2017 landslide area and areas y and z to the west (fig. 2). these areas, or subareas within them, show incoherence in all studied interferograms back to the spring of 2015. an example of this is given in fig. 2d where the last image in the pair is from just hours prior to the seismic event. the image pairs (both 6 and 12 day) spanning the seismic event have a low coherence, resulting in noisy interferograms (fig. 2e). inspection of optical images from approximately the same time as the sar images suggests that the observed loss of coherence is due to changes in snow cover over most of the slope likely due to redistribution by wind. it is thus not possible to observe this event in the interferograms as, for example, an increase in decorrelated area of the continuous active areas. instead, we have to rely on observations from the optical imagery for interpretation. this emphasises the importance of using multiple independent methods for observing landslides. possible cause of the event it is clear that the localised rock fall occurred across the time span of the shallow ml 1.9 non-tectonic seismic event. however, the interferograms spanning this event are noisy and thus we cannot yet confirm the exact cause of the event. however, by compiling all three independent datasets together, we suggest that landslide movement along a fracture could have generated the seismic signal and triggered localised rock fall near the epicentre. the landslide movement could have been either a rapid acceleration in creep of the unstable area west of the 2017 karrat landslide scar (y in fig. 2a), or the propagation of an unknown fracture related to landslide activity – potentially the westward migration of the fracture that forms the headscarp of the 2017 landslide – or both. this appears to be supported by the fact that no catastrophic mass wasting was observed in the sentinel-2 images, which might otherwise have indicated that the event was a true landslide like the 2017 karrat landslide. moreover, no large-scale calving events were observed in the two nearby glacial outlets, suggesting that glacial activity was not the cause of the seismic event. the differential interferogram prior to the event (fig. 2d) shows that areas on the slope were moving before the event, but there are no seismic signals present during this period. other similar seismic and rock fall events have been observed episodically throughout 2018 and eyewitnesses have reported activity in the area since the 2017 karrat landslide took place. the area continues to be active. outlook the approach described here has the potential to serve as a model contingency plan to gather the relevant information in the event of a landslide. this information could then be disseminated to the public and relevant authorities in the form of express reports. these methods can also be applied to quantify the temporal evolution of past landslides using archive data. sentinel-2 images are available from mid-2015 and for older events, more coarse-scale landsat images can be used back to the early seventies. sentinel-1a sar data are available from october 2014, whereas both sentinel-1a and 1b data are available from october 2016. the current glisn seismic network is available back to summer 2010 (clinton et al. 2017). the first stations were installed in greenland in 1928, and until the 1990s the network consisted of only 3–4 stations, increasing to 5–8 stations before the glisn network. prior to 2010, only very large landslides would have been observed by the seismic stations, for example the 2000 paatuut landslide (dahl-jensen et al. 2004), which also coincided with e2019430208-05 a research network station deployment (dahl-jensen et al. 2003). a similar approach could also be implimented to quantify activity in other remote areas that are prone to landslides. an obvious candidate is the south coast of the nuussuaq peninsular where several catastrophical historical and prehistorical landsides are known to have occurred (pedersen et al. 2002; dahl-jensen et al. 2004; svennevig 2019). in the future, a denser network of seismographs on a regional scale would greatly improve our capability to locate earthquakes in the area more precisely and help us to determine the cause of non-tectonic events. for example, an event that is far from a glacier terminus is less likely to be caused by glacial activity. a local seismic network around the landslide coast at karrat would further help to determine the cause of these earthquakes by pinpointing exactly where on the slope they occur relative to known structures. a similar network could also be applied to the vaigat coast where non-tectonic seismic events are also suspected to be caused by landslide activity, and where there is a historical record of landslides (pedersen et al. 2002; dahl-jensen et al. 2004). automation of the sar and optical data processing (change detection; e.g., lacroix et al. 2018) could also aid the workflow. while we are unable to precisely identify the cause of the ml 1.9 seismic event, it is clear that the combination of the various datasets is key to understanding the process involved in such events and demonstrates the benefit of a multidisciplinary approach. acknowledgments we thank the reviewers, andrée blais-stevens and erin bessette-kirton, whose comments and suggestions improved the manuscript. references bessette-kirton, e., allstadt, k., pursley, j. & godt, j. 2017: preliminary analysis of satellite imagery and seismic observations of the nuugaatsiaq landslide and tsunami, greenland. https://www.usgs.gov/ natural-hazards/landslide-hazards/science/preliminary-analysis-satellite-imagery-and-seismic?qt-science_center_objects=0#qt-science_ center_objects. clinton, j.f. et al. 2014: seismic network in greenland monitors earth and ice system. eos, transactions american geophysical union 95, 1314. https://doi.org/10.1002/2014eo020001 clinton, j., larsen, t.b., dahl-jensen, t., voss, p.h. & nettles, m. 2017: seismic observations from nuugatsiaq slide/tsunami. iris special 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precambrian rocks of the umanak area, west greenland. meddelelser dansk geologisk forening 17, 1–20. lacroix, p., bièvre, g., pathier, e., kniess, u. & jongmans, d. 2018: use of sentinel-2 images for the detection of precursory motions before landslide failures. remote sensing of environment 215. 507–516. https:// doi.org/10.1016/j.rse.2018.03.042 nettles, m. et al. 2008: step-wise changes in glacier flow speed coincide with calving and glacial earthquakes at helheim glacier, greenland. geophysical research letters 35, 1–5. https://doi. org/10.1029/2008gl036127 pedersen, s.a.s., larsen, l.m., dahl-jensen, t., jepsen, h.f., krarup, g., nielsen, t., pedersen, a.k., von platen-hallermund, f. & weng, w.l. 2002: tsunami-generating rock fall and landslide on the south coast of nuussuaq , central west greenland. geology of greenland survey bulletin 191, 73–83. porter, c. et al. 2018: arcticdem v2, https://doi.org/10.7910/dvn/ ohhukh, harvard dataverse. rosen, p.a., hensley, s., joughin, i.r., li, f.k., madsen, s.n., rodriguez, e. & goldstein, r. 2000: synthetic aperture radar interferometry. proc. of the ieee 88, 333–382. https://doi.org/10.1109/5.838084 sørensen, e.v. & guarnieri, p. 2018: remote geological mapping using 3d photogrammetry: an example from karrat, west greenland. geological survey of denmark and greenland bulletin 41, 63–66. svennevig, k. 2019: preliminary landslide mapping in greenland. geological survey of denmark and greenland bulletin 43, e2019430207. https://doi.org/10.34194/geusb-201943-02-07 how to cite svennevig, k. solgaard, a.m., salehi, s., dahl-jensen, t., merryman boncori, j.p., larsen t.b., & voss, p.h. 2019: a multidisciplinary approach to landslide monitoring in the arctic: case study of the march 2018 ml 1.9 seismic event near the karrat 2017 landslide. geological survey of denmark and greenland bulletin 43, e2019430208. https://doi.org/10.34194/ geusb-201943-02-08 *corresponding author: kristian svennevig | e-mail: ksv@geus.dk 1 geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350, copenhagen k, denmark. 2 dtu space, national space institute, technical university of denmark, ørsteds plads, 2800 lyngby, denmark. https://www.usgs.gov/natural-hazards/landslide-hazards/science/preliminary-analysis-satellite-imagery-and-seismic?qt-science_center_objects=0#qt-science_center_objects https://www.usgs.gov/natural-hazards/landslide-hazards/science/preliminary-analysis-satellite-imagery-and-seismic?qt-science_center_objects=0#qt-science_center_objects https://www.usgs.gov/natural-hazards/landslide-hazards/science/preliminary-analysis-satellite-imagery-and-seismic?qt-science_center_objects=0#qt-science_center_objects https://www.usgs.gov/natural-hazards/landslide-hazards/science/preliminary-analysis-satellite-imagery-and-seismic?qt-science_center_objects=0#qt-science_center_objects https://doi.org/10.1002/2014eo020001 https://ds.iris.edu/ds/nodes/dmc/specialevents/2017/06/22/nuugaatsiaq-greenland-landslide-and-tsunami/#seismic-observations-from-nuugatsiaq-slidetsunami https://ds.iris.edu/ds/nodes/dmc/specialevents/2017/06/22/nuugaatsiaq-greenland-landslide-and-tsunami/#seismic-observations-from-nuugatsiaq-slidetsunami https://ds.iris.edu/ds/nodes/dmc/specialevents/2017/06/22/nuugaatsiaq-greenland-landslide-and-tsunami/#seismic-observations-from-nuugatsiaq-slidetsunami https://doi.org/10.1016/s0012-821x(02)01080-4 https://doi.org/10.1016/s0012-821x(02)01080-4 https://doi.org/10.1023/b:nhaz.0000020264.70048.95 https://doi.org/10.1023/b:nhaz.0000020264.70048.95 https://doi.org/10.1126/science.1088057 https://doi.org/10.1007/s10346-017-0926-4 https://doi.org/10.1007/s10346-017-0926-4 https://doi.org/10.1016/j.rse.2018.03.042 https://doi.org/10.1016/j.rse.2018.03.042 https://doi.org/10.1029/2008gl036127 https://doi.org/10.1029/2008gl036127 https://doi.org/10.7910/dvn/ohhukh https://doi.org/10.7910/dvn/ohhukh https://doi.org/10.1109/5.838084 https://doi.org/10.34194/geusb-201943-02-07 https://doi.org/10.34194/geusb-201943-02-08 https://doi.org/10.34194/geusb-201943-02-08 mailto:rsf%40geus.dk?subject= geological survey of denmark and greenland bulletin 1, 459-526 459 late triassic – jurassic development of the danish basin and the fennoscandian border zone, southern scandinavia lars h. nielsen the continental to marine upper triassic – jurassic succession of the danish basin and the fennoscandian border zone is interpreted within a sequence stratigraphic framework, and the evolution of the depositional basin is discussed. the intracratonic permian–cenozoic danish basin was formed by late carboniferous – early permian crustal extension followed by subsidence governed primarily by thermal cooling and local faulting. the basin is separated from the stable precambrian baltic shield by the fennoscandian border zone, and is bounded by basement blocks of the ringkøbing–fyn high towards the south. in late triassic – jurassic times, the basin was part of the epeiric shallow sea that covered most of northern europe. the upper triassic – jurassic basin-fill is subdivided into two tectono-stratigraphic units by a basinwide intra-aalenian unconformity. the norian – lower aalenian succession was formed under relative tectonic tranquillity and shows an overall layer-cake geometry, except for areas with local faults and salt movements. deposition was initiated by a norian transgression that led to shallow marine deposition and was accompanied by a gradual climatic change to more humid conditions. extensive sheets of shoreface sand and associated paralic sediments were deposited during short-lived forced regressions in rhaetian time. a stepwise deepening and development of fully marine conditions followed in the hettangian – early sinemurian. thick uniform basinwide mud blankets were deposited on an open storm-influenced shelf, while sand was trapped at the basin margins. this depositional pattern continued until late toarcian – early aalenian times when the basin became restricted due to renewed uplift of the ringkøbing–fyn high. in middle aalenian – bathonian times, the former basin area was subjected to deep erosion, and deposition became restricted to the fault-bounded sorgenfrei–tornquist zone. eventually the fault margins were overstepped, and paralic–marine deposition gradually resumed in most of the basin in late jurassic time. thus, the facies architecture of the norian – lower aalenian succession reflects eustatic or large-scale regional sea-level changes, whereas the middle aalenian – volgian succession reflects a strong tectonic control that gradually gave way to more widespread and sea-level controlled sedimentation. the uplift of the ringkøbing–fyn high and most of the danish basin occurred concurrently with the uplift of the north sea and a wide irregular uplifted area was formed, which differs significantly from the postulated domal pattern. keywords: danish basin, fennoscandian border zone, upper triassic – jurassic, sedimentology, sequence stratigraphy, basin development, intra-aalenian unconformity, sea-level control geological survey of denmark and greenland, geocenter copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: lhn@geus.dk geological survey of denmark and greenland bulletin 1, 459–526 (2003) © geus, 2003 the deeply buried upper triassic – jurassic succession of the danish basin and the fennoscandian border zone has been a target for exploration activities since 1935 and c. 60 deep wells have been drilled for hydrocarbons, geothermal energy or gas storage in the danish area (sorgenfrei & buch 1964; nielsen & japsen 1991). outcrops only occur along the basin margin in skåne and bornholm (gravesen et al. 1982, norling et al. 1993). the aims of the paper are to interpret the upper triassic – jurassic basin-fill within a sequence stratigraphic framework, to describe the spatial distribution of facies and depositional sequences, and to provide a scheme for comparison of the deeply buried basinal successions with the basin margin successions partially exposed in skåne and on bornholm. the results are presented in chronostratigraphic schemes. a relative sealevel curve is constructed and compared to eustatic curves and curves from other parts of the jurassic epeiric sea, and subsidence curves for five selected positions in the basin are presented. the formation and timing of the significant intra-aalenian unconformity is discussed and compared to the coeval north sea unconformity. the paper concludes with a description of the late triassic – jurassic evolution of the basin. the study utilises well logs from more than 40 wells and a large number of geographically and stratigraphically widely distributed cores from the gassum, fjerritslev, haldager sand and flyvbjerg formations. facies associations are defined from sedimentological core logs, because associations rather than individual facies can be recognised on well logs. non-cored sections are interpreted by means of well logs and information from sidewall cores, cuttings samples, biostratigraphic and palaeoecological data. tectonic setting the danish basin is roughly synonymous with the danish embayment (sorgenfrei & buch 1964; larsen 1966; michelsen 1975), the danish part of the norwegian–danish basin (bertelsen 1978; michelsen 1978) and the danish subbasin (michelsen 1989a, b). the basin is an intracratonic, permian–cenozoic structure that trends wnw–ese. it is bounded by basement blocks of the ringkøbing–fyn high to the south and by the fennoscandian border zone to the north-east (figs 1, 2). the border zone demarcates the transition to the stable precambrian baltic shield and includes the sorgenfrei–tornquist zone and the skagerrak– kattegat platform (sorgenfrei & buch 1964; bergström 1984; eugeno-s working group 1988). the sorgenfrei– tornquist zone forms the northern segment of the tornquist zone which is a long-lived fundamental tectonic feature. it converges with the teisseyre–tornquist zone via the rønne graben offshore bornholm (fig. 1), where the danish basin passes into the polish trough. the sorgenfrei–tornquist zone is strongly block-faulted, 30–50 km wide, with tilted palaeozoic fault blocks unconformably overlain by thick mesozoic deposits that show pronounced late cretaceous – early cenozoic tectonic inversion. the skagerrak–kattegat platform is a stable area to the north-east where the mesozoic deposits onlap lower permian, lower palaeozoic and precambrian crystalline rocks in tilted fault blocks and gradually thin out towards the baltic shield (fig. 2). the deepest regional surface mappable by reflection seismic data in the danish basin and fennoscandian border zone is the top pre-zechstein surface, which is a pronounced unconformity truncating tilted fault blocks in most of the area (vejbæk 1989, 1997; britze & japsen 1991; michelsen & nielsen 1991, 1993; vejbæk & britze 1994). the unconformity is penetrated by wells which show the occurrence of precambrian crystalline rocks on the ringkøbing–fyn high (glamsbjerg-1, grindsted-1, ibenholt-1 and jelling-1) and the skagerrak– kattegat platform (frederikshavn-1) and lower palaeozoic sedimentary rocks in the danish basin and fennoscandian border zone (nøvling-1, rønde-1, slagelse-1 and terne-1; sorgenfrei & buch 1964; poulsen 1969, 1974; christensen 1971, 1973; larsen 1971, 1972; michelsen & nielsen 1991, 1993; nielsen & japsen 1991). the wells cut the unconformity on footwall blocks or on hangingwall block crests, where deep erosion has occurred, thus making accurate dating of the rifting impossible. in contrast, the hans-1 and sæby-1 wells are located in the deep hangingwalls of tilted fault blocks close to the footwall fault (michelsen & nielsen 1991, 1993). in sæby-1, on the skagerrak–kattegat platform, the unconformity separates triassic sediments from syn-rift rotliegendes volcaniclastic rocks. in the sorgenfrei–tornquist zone, hans-1 penetrates a prerift succession of clastic sediments and extrusive volcanic rocks, of presumed late carboniferous age (fig. 3). the succession is unconformably overlain by a thick rotliegendes syn-rift prism of alluvial conglomerates and sandstones and lacustrine mudstones. the syn-rift succession is overlain by marginal to non-marine zechstein deposits. in the nearby terne-1 well, upper carboniferous intrusive volcanic rocks occur, and the volcanic rocks in hans-1 and terne-1 seem to be roughly 460 461 6° e 10 °e 14 °e 57 °n 58 °n 56 °n 55 °n n or w ay sk ån e g er m an y t he n et he rl an ds sw ed en bo rn ho lm h or n g ra be n h an ö ba y st en lil le w el ls u lle rs le v1 h or se ns -1 vo ld um -1 r øn de -1 te rn e1 g as su m -1 h ob ro -1 ve m b1m ej ru p1 r ød di ng -1 fa rs ø1 å rs -1 h yl le bj er g1 ve ds te d1 fl yv bj er g1 sæ by -1 t hi st ed -1 , 4 t hi st ed -3 n øv lin g1 g ri nd st ed -1 m or s1 o dd es un d1 sk iv e1 k vo ls -1 bø rg lu m -1 h al da ge r1 sk ag en -2 fr ed er ik sh av n1, 2, 3 g la m sb je rg -1 sl ag el se -1 r ød by -1 la vø -1 h an s1 fj er ri ts le v1, -2 f1 in ez -1 k -1 fe lic ia -1 je lli ng -1 a nh ol t4 d an is h c en tr al g ra be n w el l fa ul t n at io na l b or de r h ig h 10 0 km st ru ct ur al e le m en ts o f so ut he rn s ca nd in av ia so rg en fre i– to rn qu ist z on e bø rg lu m fa ult h ald ag er fa ult fje r r its le v fa ul t d an ish ba sin sk ur up h ig h ri ng kø bi ng –f yn h ig h sk ag er ra k– h i. g r. fj er rit sle v tr ou gh k at te ga t pl at fo rm r øn ne g r. ea st n or th s ea h ig h tt z fi g. 1 . m ap s h o w in g th e w el l l o ca tio n s an d th e p ri n ci p al s tr u ct u ra l u n its o f th e st u d y ar ea ; t h e o u tli n e o f th e r in gk ø b in g– fy n h ig h is b as ed o n th e d is tr ib u tio n o f z ec h st ei n d ep o si ts . t h e so rg en fr ei –t o rn q u is t z o n e m er ge s w ith t h e te is se yr e– to rn q u is t z o n e (t t z ) o ff sh o re b o rn h o lm i n t h e r ø n n e g ra b en . t h e d as h ed r ed l in e in d ic at es t h e tr an se ct c o ve re d b y th e ge o se ct io n i n f ig . 2. m o d if ie d f ro m m ic h el se n & n ie ls en ( 19 91 ) an d v ej b æ k (1 99 7) . contemporaneous with the earliest volcanic rocks in the oslo graben and dolerite dykes in southern sweden (bergström et al. 1982; ro et al. 1990). the principal phase of rifting of the danish basin and the fennoscandian border zone thus occurred in the late carboniferous – early permian at the same time as, or slightly later than, rifting of the oslo graben (ro et al. 1990; michelsen & nielsen 1991, 1993). the tilted fault block crests are deeply truncated by the mid-permian unconformity showing that regional post-rift thermal subsidence was somewhat delayed (vejbæk 1997). the unconformity defines the base of the post-rift succession and is overlain by a relatively complete succession of upper permian, mesozoic and cenozoic deposits that is c. 5–6.5 km thick along the basin axis and more than 9 km locally in the sorgenfrei–tornquist zone and the himmerland graben (fig. 2). isochore maps of the triassic and jurassic – lower cretaceous successions show a relatively uniform regional thickness over most of the basin except for areas influenced by local halokinetic movements, indicating relatively uniform thermal subsidence (vejbæk 1989, 1997; britze & japsen 1991; japsen & langtofte 1991). although the thick upper permian – triassic succession indicates rapid subsidence that exceeds normal thermal contraction, a prolonged or new rifting phase is precluded by the general lack of pronounced extensional faulting in the mesozoic succession (vejbæk 1989, 1997). the evaporitic and continental facies show that the basin was never under-filled, and phase transformations in the deep crust have been proposed to explain the rapid early post-rift subsidence (vejbæk 1989). the great thicknesses of the mesozoic in the himmerland graben and the fjerritslev trough were facilitated by transtensional strike-slip movements in the sorgenfrei– tornquist zone and large-scale salt movements (pegrum 1984; vejbæk 1989; christensen & korstgård 1994; mogensen 1996). a general shallowing of the basin towards the ringkøbing–fyn high is indicated by thinning of the 462 nøvling-1 ringkøbing–fyn high danish basin haldager-1 flyvbjerg-1 frederikshavn-1 fjerritslev trough sorgenfrei–tornquist zone sw ne sw ne skagerrak–kattegat platform 10 km 0 1 2 3 4 t w t ( se c) 0 1 2 3 4 t w t ( se c) zechstein middle–lower jurassic lower–middle triassic cenozoic upper cretaceous lower cretaceous/ upper jurassic upper triassic rotliegendes basement palaeozoic fig. 2. a regional sw–ne geosection through the danish basin and the fennoscandian border zone. for location, see fig. 1. modified from vejbæk (1990, 1997). twt, two-way travel time. zechstein – lower jurassic and upper jurassic – lower cretaceous successions although erosion at the base of the middle jurassic and the cretaceous have obscured the original distribution of the triassic – lower jurassic on the high. the high was probably formed at the same time as the danish basin as an area of less stretching (vejbæk 1997). marginal facies were developed along the high in late permian time and it probably formed a barrier between the southern and northern zechstein basins (ziegler 1982; stemmerik et al. 1987). in late triassic – early jurassic times, the high became flooded during periods of high sea level (michelsen 1975, 1978; bertelsen 1978; l.h. nielsen, l. hamberg and e.b. koppelhus in: nielsen 1995). in middle jurassic time, it was uplifted causing a major change in the basin configuration (michelsen 1978; koch 1983; nielsen 1993, 1995). upper triassic – jurassic stratigraphy the upper triassic – jurassic succession is more than 1 km thick and includes the skagerrak, vinding, gassum, fjerritslev, haldager sand, flyvbjerg, børglum and frederikshavn formations (fig. 4). the stratigraphic scheme currently applied is based on sorgenfrei & buch (1964), larsen (1966), michelsen (1975, 1978, 1989a), bertelsen (1978, 1980) and michelsen et al. (2003, this volume). claystones, marls and oolitic carbonates of the norian vinding formation were deposited in a shallow, restricted marine environment in central parts of the basin and form a succession, 40–90 m thick, and locally up to 200 m (bertelsen 1978, 1980; nielsen & japsen 1991). concurrently, alluvial arkosic sandstones and lacustrine claystones of the skagerrak formation were deposited in the northern and north-eastern, marginal parts of the 463 basementbasement lower palaeozoic triassic triassic triassic jurassic jurassic jurassic cambrian cambrian l. cretaceous? ordovician-silurian ordovician-silurian rotliegendeu.car.u.car. u. carboni-ferous børglum fault zechstein 0 1 2 3 0 1 2 3 t w t ( se c. ) t w t ( se c. ) z z r r 5 km ? ? hans-1 l. cretaceous? jurassic triassic zechstein rotliegendes u. carboniferous ordovician–silurian cambrian basement sw ne sorgenfrei–tornquist zone fig. 3. a sw–ne geosection through the sorgenfrei–tornquist zone, intersecting the hans-1 well (see fig. 1). note the tilted palaeozoic fault blocks, the rotliegendes syn-rift prisms and the regional mid-permian unconformity overlain by the relatively undisturbed upper permian – mesozoic basin-fill. modified from michelsen & nielsen (1991, 1993). twt, two-way travel time. 464 basin. the uppermost norian – lower sinemurian gassum formation overlies the skagerrak formation along the basin margin, and interfingers with the upper part of the vinding formation in the deep part of the basin (bertelsen 1978; l.h. nielsen, l. hamberg and e.b. koppelhus in: nielsen 1995). the thickness varies from 50–150 m in the central part of the basin to more than 300 m locally in the sorgenfrei–tornquist zone, and the formation consists of interbedded fineto mediumgrained, occasionally coarse-grained and pebbly sandstones, heteroliths, mudstones and few thin coaly beds (bertelsen 1978; michelsen & nielsen 1991). bivalves, foraminifers and rare ammonites occur in addition to common spores, pollen and dinoflagellate cysts. an overall fluvio-deltaic, deltaic to tidally-influenced shallow marine environment has been proposed (larsen 1966; bertelsen 1978; nielsen et al. 1989). the presence of several regressive shoreface sandstones of wide lateral extent shows, however, that the formation was formed under the influence of repeated sea-level fluctuations and not by simple deltaic progradation (hamberg et al. 1992; hamberg 1994; l.h. nielsen, l. hamberg and e.b. koppelhus in: nielsen 1995; hamberg & nielsen 2000). the overlying lower jurassic fjerritslev formation is dominated by marine claystones and mudstones containing ammonites, bivalves, foraminifers, ostracods and dinoflagellate cysts (nørvang 1957; larsen 1966; michelsen 1975; pedersen 1986; dybkjær 1988, 1991; poulsen 1996). the formation shows very variable thickness due u m l u l u l u m l u l u l annero fm a nn er o fm m ar ie da l f m vilhelmsfält fm fyledalen clay fortuna marl glass sand mb fuglunda mb r ya f m h ög an äs f m helsingborg mb döshult mb pankarp mb katslösa mb rydebäck mb röddinge fm ? sorthat fm bagå fm hasle fm galgeløkke mb sose bugt mb munkerup mb risebæk mb r øn ne f m gassum fm fjerritslev fm haldager sand fm hettangian sinemurian pliensbachian toarcian aalenian bajocian bathonian callovian oxfordian m id dl e u pp er lo w er u pp er ju ra ss ic tr ia ss ic sw ne nw se flyvbjerg fm ? ?? ? nytorp sand kågeröd fm k åg er öd fm vallåkra mb bjuv mb oddesund fm børglum fm frederikshavn fm (part) norian rhaetian kimmeridgian volgian h öö r ss t. vinding fm skagerrak fm system stage danish basin bornholm onshoreskåne fig. 4. lithostratigraphic scheme for the upper triassic – jurassic of the danish basin and the fennoscandian border zone based on bertelsen (1978, 1980), michelsen (1978, 1989a), gravesen et al. (1982), sivhed (1984), norling et al. (1993), koppelhus & nielsen (1994), l.h. nielsen, l. hamberg and e.b. koppelhus (in: nielsen 1995), ahlberg et al. (2003, this volume) and michelsen et al. (2003, this volume). 465 to erosional truncation in the south-western and central parts of the basin and above salt structures, and a maximum thickness of more than 1000 m is reached in the fault-bounded fjerritslev trough (fig. 2). lower jurassic mudstones and sandy mudstones are also present south of the ringkøbing–fyn high (bertelsen & michelsen 1970; michelsen 1973, 1975). the formation covers the early jurassic time interval, and also includes mudstones of latest rhaetian and early aalenian age (dybkjær 1991; michelsen & nielsen 1991; poulsen 1996). the transition from the gassum formation to the fjerritslev formation occurred in several steps ranging from latest rhaetian in the central parts of the basin to early sinemurian at the north-eastern margin reflecting the overall early jurassic eustatic sea-level rise (bertelsen 1978; michelsen 1978, 1989b; hallam 1988; dybkjær 1991; michelsen & nielsen 1991; nielsen 1995). the haldager sand formation erosionally overlies the fjerritslev formation and consists of fineto very coarsegrained, occasionally pebbly sandstones, siltstones, mudstones and coaly beds (michelsen 1978, 1989a; koch 1983). koch (1983) proposed a general braided fluvial to deltaic depositional environment. the formation is absent on and along the ringkøbing–fyn high and is thin and patchy in large parts of the basin except for in rim-synclines close to salt structures (e.g. mors-1, thisted-3; fig. 1). in the south-western part of the basin, it consists of fineto coarse-grained fluvial sandstone units, 1–10 m thick, and thickens to more than 150 m towards the north-east in the sorgenfrei–tornquist zone, where the formation includes paralic and shallow marine sandstones and mudstones. dating of the formation is generally poor giving a broad middle jurassic age (michelsen 1978, 1989a; michelsen & nielsen 1991; poulsen 1992a, 1996). it is overlain by transgressive paralic–marine mudstones and sandstones of the oxfordian flyvbjerg formation, which shows roughly the same distribution as the haldager sand formation. expansion and deepening of the basin is reflected by the change to the overlying kimmeridgian–ryazanian marine mudstones of the børglum formation and volgian–ryazanian marine to paralic siltstones, sandstones and mudstones of the frederikshavn formation that overstep the limits of the flyvbjerg formation (michelsen 1978, 1989a; michelsen & nielsen 1991; poulsen 1996). sedimentology of the gassum, fjerritslev and haldager sand formations the upper triassic – middle jurassic siliciclastic deposits are composed of a number of relatively uniform and recurrent continental, paralic, nearshore and offshore facies that are grouped into six associations (l.h. nielsen, l. hamberg and e.b. koppelhus in: nielsen 1995; hamberg & nielsen 2000). the typical well-log pattern for each association is described by comparing cores and well logs (table 1). the logs show sharp-based, blocky motifs with fairly consistently low gamma-ray readings, increasing in the upper part. the sp-log pattern is very blocky in old wells. the gamma-ray readings are generally lower than those of the shoreface sandstones, although atypically high values may be recorded in zones with abundant claystone clasts and/or diagenetic clay minerals. sp-logs from the uppermost part of the skagerrak formation and the gassum and haldager sand formations show uniform rightwards deflections. the change from variegated (skagerrak or lower gassum formation) to greyish claystones (gassum formation) on the skagerrak–kattegat platform is marked by a significant rightwards deflection. the logs show sharp-based motifs, the basal part yielding low gamma-ray values followed by an overall increase in values upwards reflecting the general fining-upwards of the association. bellor funnel-shaped motifs occur internally. discrimination between the lower sandy part of estuarine and fluvial deposits by log motifs alone may be difficult, but the heterolithic component of the estuarine deposits generally causes more variable log patterns. the log pattern is very variable. some facies show an increase upwards in gamma-ray values in the basal part, reflecting increasing mudstone content, followed by a decrease reflecting the incoming of sandstones. others show several decreasing-upwards trends reflecting coarseningand cleaning-upwards heteroliths. a distinct decrease in gamma-ray values at the top marks a marine transgressive erosion surface overlain by sandstones. wave-reworked sandstones or channel sandstone units within the background heterolith facies are indicated by distinct gamma-ray lows. a serrated log pattern indicates a dominance of stacked, fining-upwards channel units (see figs 7a, 9b). the typical gamma-ray and sp-log motifs are sharp-based and smooth, blocky to funnel-shaped, reflecting several almost amalgamated, coarsening-upwards sandstone units or two to three sandstone units separated by heteroliths. the general coarsening-upwards trend shown by cores is reflected by a weak decreasing-upwards gamma-ray trend or leftwards deflection of the sp-log. this trend and commonly slightly higher gamma-ray values serves to discriminate shoreface from fluvial sandstones. the well-log motifs of this facies association show high to intermediate gamma-ray readings, corresponding sp-values and low sonic velocities forming both relatively uniform units and units showing decreasingor increasing-upwards signals (see figs 8, 10, 13). fluvial facies association lacustrine facies association estuarine facies association lagoonal facies association shoreface and foreshore facies association marine offshore facies association table 1. typical well-log patterns of the facies associations 466 core 18 sp-log lagoonal flooding surface/ts 10 marine flooding surface transgressive systems tract lowstand systems tract clay si sand 2 4 mm shoreface soil horizon soil horizon fluvial channel fluvial channel – – – – sp-log lowstand systems tract transgressive systems tract flooding surface lower shoreface to offshore lower shoreface tsme/ts 12 børglum-1 core 17 3 m 2 1 0 4 m 3 2 1 0 1 m 0 core 19 highstand systems tract lowstand systems tract sb 10 sp-log shoreface – offshore transition zone fluvial channel shoreface core 20 clay si sand 2 4 mm ✩ ✩ ✩ fs tsme sp-log highstand systems tract offshore shoreface transgressive sand fluvial– estuarine 4 m 3 2 1 0 ✩ core 21 clay si sand 2 4 mm lowstand systems tract transgressive systems tract ts 9 fluvial? beach 4 m 3 2 1 0 clay si sand 2 4 mm clay si sand 2 4 mm – 467 fluvial facies association in the gassum formation, the facies association is most common in the fjerritslev trough, but it also occurs in the basin centre. it was cored in the børglum-1, flyvbjerg-1, frederikshavn-2, horsens-1, thisted-3, vedsted-1 and års-1 wells (figs 5, 6, 7, 8, 9a, 12, 15). the association dominates the haldager sand formation and was cored in the farsø-1, frederikshavn-1 and -3, haldager-1, skagen-2, vedsted-1 and års-1 wells (fig. 10b). the association typically consists of fining-upwards units, 2–14 m thick (most commonly 4–6 m), beginning with a sharp, erosional base overlain by sandstones that contain clasts of claystone, coal, coalified wood and occasional pebbles at the base (figs 9a, 10b, 11a, b). the sandstones, which fine weakly upwards, are overlithology mudstone siltstone sandstone pebble lags coal clay ironstone concretions calcareous cement pyritic concretions glauconite carbonaceous detritus mudstone chips clast clast clast clast biogenic structures degree of bioturbation cryptobioturbation by amphipods chondrites isp. diplocraterion isp. helminthopsis isp. rhizocorallium isp. skolithos isp. teichichnus isp. thalassinoides isp. zoophycos isp. horizontal burrows large (crustacean?) burrows escape trace rootlets sedimentary structures erosional surface parallel bedding/lamination planar cross-bedding trough cross-bedding low-angle cross-bedding hummocky cross-stratification small-scale hummocky cross-stratification gutter casts cross-lamination and climbing ripples wave ripples flaser and wavy bedding lenticular and silt-streaked bedding disturbed bedding load structures water escape structures synaeresis cracks desiccation cracks fossils gastropod bivalve fragments plant wood stem ammonite belemnite facing page: fig. 5. core logs of the gassum (cores 18–21) and fjerritslev formations (core 17) in the børglum-1 well (for location of the cores in the well, see fig. 20). core 21 shows a fluvial channel unit belonging to the lowstand systems tract (lst) of the fj 1 sequence. it is capped by transgressive surface ts 9 overlain by a weakly bioturbated beach sandstone of the transgressive systems tract (tst). no well logs were obtained for this interval. cores 20 and 19 show the highly variable deposits of the highstand systems tract (hst) of sequence fj 1. uppermost in core 19 is a coarse-grained, cross-bedded erosionally-based fluvial sandstone, the base of which marks sb 10 and the base of the fj 2 sequence. core 18 shows fluvial channel units with roots and soil horizons of the lst capped by ts 10 (sequence fj 2); ts 10 is overlain by thinly preserved lagoonal deposits below a marine shoreface sandstone. core 17 shows lower shoreface sandy heteroliths (lst; sequence fj 4) overlain by transgressive, deeper water deposits (tst). the flooding surface ts 12 is defined by a lamina of very coarse-grained sand draped by 2–5 cm of cemented mudstone. core depth is corrected to log depth by subtracting 3 ft. the accompanying legend (above) is applicable to all the core logs (figs 5–14). fs, flooding surface; sb, sequence boundary; ts, transgressive surface; tsme, transgressive surface of marine erosion. horsens-1 core 4 sp-log lowstand systems tract transgressive systems tract tsme/ts 4 shoreface shoreface fluvial channel clay si sand 2 4 mm 2 m 0 1 fig. 6. core log of the gassum formation in the horsens-1 well. fluvial sandstones (lst) are overlain by transgressive shoreface sandstones (tst) of the fourth-order sequence 4, sequence vi 1 (core depths are corrected to log depths by subtracting c. 10 m; the slight misfit between core and sp-log is caused by poor recovery of the interbedded offshore mudstones). 468 lain by highly carbonaceous claystones and coaly beds with roots, or are cut by erosion surfaces overlain by a similar fluvial unit, lacustrine or marine deposits. palynological samples yield only poor assemblages of spores and pollen (koch 1983; e.b. koppelhus and n.e. poulsen, personal communications 1994; l.h. nielsen, l. hamberg and e.b. koppelhus in: nielsen 1995). the sandstones are mostly mediumto coarsegrained, occasionally fine-grained, micaceous, carbonaceous, and moderately sorted with subangular to subrounded quartz grains. the colour is typically greyish, occasionally yellowish or brownish. coarse-grained sand with small quartz pebbles occurs locally in the haldager sand formation. comminuted organic debris and coalified fragments of leaves, stems and wood occur commonly in both the haldager sand and gassum sb 5 sb 4 mfs 3 sb 3 thisted-3 cores 4, 5 highstand systems tract transgressive systems tract lowstand systems tract lowstand systems tract transgressive systems tract lagoonal flooding surface 1 0 fluvial channel shoreface lagoon lagoon swamp fluvial channel lagoon with tidal creeks lagoon offshore clay si sand 2 4 mm 15 m 10 5 3 tst gamma-ray a this and facing page: fig. 7. core logs of the gassum formation in the thisted-3 well. cores 4 and 5 (a) show lagoonal and fluvial deposits of the fourth-order sequences 3 and 4, belonging to the hst of the vi 1 sequence. vi 1 is bounded above by sb 5 which is overlain by a lst of fluvial sandstones (a; uppermost core 4) and fluvial–estuarine sandstones with minor mudstones of the fourth-order sequence 5 (b; core 3) of the ga 1 sequence. c: core 2 exhibits lagoonal deposits of the tst overlain by a thin hst and shoreface sandstones of the forced regressive systems tract (frst) of sequence 5, ga 1 sequence. the upper part of core 2 shows shoreface sandstones (lst) overlain by shoreface to offshore transition zone sandstones and heteroliths of sequence 6, ga 1 sequence. for legend, see fig. 5. fluvial– estuarine 8 m 6 4 2 1 7 5 3 0 lowstand systems tract gamma-ray thisted-3 core 3 clay si sand 2 4 mm b 469 thisted-3 core 2 gamma-ray lower shoreface to offshore transition zone lagoon lagoon lagoon upper shoreface 18 m 15 10 5 0 transgressive systems tract transgressive systems tract lowstand systems tract forced regressive systems tract hst rsme mfs 5 mfs 6 tsme/ts 6 sb 6 clay si sand 2 4 mm c 470 clay si sandsp-log 3 m 2 1 0 forced regressive systems tract vedsted-1 core 7 0 1 2 3 m core 9 lagoonal flooding surface clay si sandsp-log lowstand systems tract core 10 sb 3 clay si sand 5 m 4 3 2 1 0 sp-log lowstand systems tract highstand systems tract 2 4 mm shoreface to foreshore offshore – lower shoreface foreshore soil horizon lagoon fluvial– estuarine channel fluvial channel lake 2 4 mm 2 4 mm formations. the sandstones show trough and planar cross-bedding, parallel lamination and cross-lamination. the trough cross-bedded sets are typically 0.05–0.4 m thick and form cosets 1–4 m thick. the planar cross-bedded sets are 0.1–1.5 m thick, and may be isolated or occur in cosets commonly interbedded with trough cross-bedded or parallel-laminated sandstones. fine-grained, crosslaminated sandstones occur rarely in both formations. the coarse-grained, pebbly, parallel-laminated beds indicate upper flow regime conditions, while the trough and planar cross-bedding were formed by migrating sinuous and straight-crested dunes under strong currents of the lower flow regime (harms et al. 1982; allen 1984). the thick, planar cross-bedded sets represent large bars (cant & walker 1978). the cross-lamination was formed by small ripples driven by weak currents. the fining-upwards units that overlie sharp, erosional bases and consist of cross-bedded cosets and parallellaminated beds, capped by carbonaceous claystones and coal seams, indicate deposition in sandy fluvial channels. this is supported by the occurrence of nonmarine palynomorphs, the absence of marine indicators, and the presence of large plant fragments. the generally high content of dispersed organic debris indicates vegetated interfluve areas. the coarse grain size, the poorly developed fining-upwards trends and the general scarcity of small-scale structures and fine-grained material suggest deposition in braided streams rather than meandering or anastomosing channels (allen 1965; miall 1977; smith 1983). the log-patterns indicate that several thin channel units amalgamate to form fluvial sandstone units up to 25 m thick in both the gassum and haldager sand formations in the fjerritslev trough and up to 12 m thick units on the skagerrak–kattegat platform. lacustrine facies association in the gassum formation, this association is primarily identified in the flyvbjerg-1 and vedsted-1 wells in the fjerritslev trough (fig. 12), but was also cored in gassum-1 (fig. 13c). it is uncommon in the haldager sand formation where it was cored in the farsø-1 and skagen-2 wells (fig. 10b). in the gassum formation, the association primarily consists of massive to poorly laminated, micaceous and weakly carbonaceous, silty and dark grey mudstones forming units 7–26 m thick. variegated mudstones occur in the lower part of the formation. similar mudstones are present in the uppermost part of the skagerrak formation (frederikshavn-1 and -2, skagen-2, vedsted-1; fig. 8). in the haldager sand formation, the association consists of grey to dark grey siltstones and mudstones with parallel lamination, lenticular bedding and scattered cross-lamination. the palynomorph assemblages are very poor and non-marine (bertelsen 1978; koch 1983; l.h. nielsen, l. hamberg and e.b. koppelhus in: nielsen 1995). the structures and the fine grain size indicate deposition mainly from suspension, and the palynomorphs indicate freshwater conditions. the change from variegated claystones to greyish and more carbonaceous claystones reflects the general change in climate from hot and arid in the triassic to warm and humid in the jurassic. estuarine channel facies association this facies association is common in the gassum formation, where it was cored in the børglum-1, gassum-1, stenlille-1 and -6, thisted-3 and vedsted-1 wells (figs 5, 7, 8, 14). it consists of erosionally-based, fining-upwards units, 5–25 m thick, locally capped by coal or heteroliths with rootlets. cores from the børglum, thisted and vedsted wells mainly show cross-bedded, fineto medium-grained, occasionally coarse-grained sandstones. they commonly contain abundant mudstone clasts, comminuted organic debris and fragments of leaves and stems, probably of the genus equisetites. thin mudstone drapes and ripple-foreset laminae occur in places. the palynomorph assemblages contain the alga botryococcus sp., the dinoflagellate dapcodinium priscum and the acritarch micrhystridium sp. (l.h. nielsen, l. hamberg and e.b. koppelhus in: nielsen 1995). the thickest and best developed succession is from stenlille-1 where cores show fine-grained, cross-bedded 471 facing page: fig. 8. core logs of the skagerrak and gassum formations in the vedsted-1 well. core 10, within the vi 1 sequence (skagerrak formation) shows lacustrine mudstones (fourthorder sequence 2) erosionally overlain by fluvial sandstones that are very coarse-grained and pebbly with rock fragments, weathered feldspars and claystone clasts (lst of sequence 3). core 9 (ga 1 sequence) shows fluvial–estuarine and lagoonal deposits of the late lst of the fourth-order sequence 5 (core depths are corrected to log depths by subtracting c. 6 m). core 7 shows parts of two forestepping parasequences with marine sandstones and mudstones belonging to the frst of sequence fj 2 (core depth is corrected to log depth by subtracting c. 5 m; slight misfit between core and log signal is caused by poor core recovery). for location of the vedsted-1 cores, refer to figs 19, 20; for legend, see fig. 5. 472 års-1 core 8 core 9 lowstand systems tract sb 5 lowstand systems tract highstand systems tract sb 4 offshore mfs 3 lagoonal flooding surface gamma-ray 10 m 5 0 0 5 8 m lagoon fluvial channels fluvial braided channels gamma-ray clay si sand 2 4 mm clay si sand 2 4 mm a 473 års-1 core 6 transgressive systems tract lagoon 0 5 9 m tidal creek tidal creek core 7 ✩ ✩ ✩ regressive surface of marine erosion mfs 5 tsme transgressive systems tract highstand systems tract 5 m 0 upper shoreface offshore offshore gamma-ray gamma-ray clay si sand 2 4 mm clay si sand 2 4 mm b this and facing page: fig. 9. core logs of the gassum formation in the års-1 well. a: core 9, within the vi 1 third-order sequence, shows marine mudstones and sandstones (hst) of the fourth-order sequence 3 erosionally overlain by cross-bedded, coarse-grained fluvial sandstones with abundant mudstone clasts and organic debris at the base constituting the lst of the fourth-order sequence 4 (see fig. 11a). core 8 shows fluvial sandstones of the fourth-order sequence 5 (ga 1 sequence) overlain by transgressive lagoonal deposits. b: core 7 shows marine deposits of sequence 5 (ga 1 sequence). note the well-defined backstepping and forestepping pattern defining mfs 5, and a sharply developed transgressive surface of marine erosion (shown in fig. 11d). core 6 illustrates lagoonal mudstones and sandstones overlying estuarine deposits (below core) of the late lst of sequence fj 1. for legend, see fig. 5. 474 farsø-1 cores 3, 4 highstand systems tract highstand systems tract transgressive systems tract mfs 3 rsme clay si sand 2 4 mm shoreface offshore – shoreface transition zone offshore offshore offshore lower shoreface lower shoreface foreshore shoreface tsme/sb 3 0 5 10 15 20 25 30 35 m tsme/sb 4 gamma-ray from sidetrack forced regressive systems tract a 475 lagoonal transgressive surface braided river lacustrine flooding surface sb (19–20) 21 offshore cl. si sand 24 mm lowstand systems tract gamma-ray lagoon transgressive systems tract lowstand systems tract braided river braided river lake transgressive systems tract farsø-1 cores 1, 2 m haldager sand fm 25 20 15 5 0 sb 22 ts 22 ts 21 flyvbjerg fm fjerritslev fm b this and facing page: fig. 10. core logs of the gassum, fjerritslev, haldager sand and flyvbjerg formations in the farsø-1 well. a: cores 3 and 4 show marine shoreface to offshore deposits of the tst, hst and frst of the fourth-order sequence 3 of the vi 1 sequence (see fig. 19). a core photo of the base of the frst is shown in fig. 11c. b: cores 1 and 2 show offshore marine mudstones of sequence fj 9 erosionally overlain by fluvial sandstones (lst) and lacustrine mudstones (tst) of sequence ha 3 (see fig. 24). the base of the fluvial sandstones marks the ‘base middle jurassic unconformity’ (see core photo in fig. 11b). the upper part of the core shows fluvial sandstones (lst) and transgressive lagoonal deposits (tst) of sequence fl 1. modified from koch (1983). for legend, see fig. 5. 476 5 cm 5 cm 5 cm 5 cm 5 cm 5 cm5 cm 5 cm a i j k l b c d fj er ri ts le v fm h al da ge r sa nd f m o ffs ho re rsme tsme br ai de d flu vi al 477 5 cm 5 cm 5 cm 5 cm 5 cm h m e f g fig. 11. core photos exhibiting the main facies of the upper triassic – jurassic in the danish basin. a: coarse-grained fluvial sandstones with abundant mudstone clasts and organic debris; the base of this core piece is an erosional surface defining sb 4. års-1, core 9, gassum formation (fig. 9a). b: cross-bedded fluvial sandstones of the haldager sand formation erosionally overlying marine mudstones of sequence fj 9, (fjerritslev formation, f-iv member) marking the ‘base middle jurassic unconformity’, consisting of the amalgamated sequence boundaries sb 19–21. farsø-1, core 2 (see figs 10, 24). c: hummocky cross-stratified shoreface sandstones with siderite-cemented clasts sharply overlying offshore bioturbated silty mudstones; the boundary represents a regressive surface of marine erosion (rsme) at the base of the frst of the fourth-order sequence 3 (vi 1 thirdorder sequence). farsø-1, core 3, gassum formation (fig. 10a, 23.2 m). d: coarse-grained, cross-bedded shoreface sandstones with abundant pyrite aggregates, capped sharply by a transgressive surface of marine erosion (tsme) overlain by burrowed sandstones (teichichnus) with swaley cross-stratification, sandy fossiliferous heteroliths with waveripple lamination and mudstones. års-1, core 7, gassum formation (fig. 9b, 1.2 m). e: the lowermost 2–3 cm of the illustrated core consists of shoreface sandstones of the hst of sequence 2, which is erosionally capped by tsme/sb 3 overlain by a thin veneer of coarsegrained sandstone and laminated and wave-rippled, fine-grained sandstones; these pass upwards into wavy and lenticular bedded mudstones with laminae and lenses of finegrained sandstones. farsø-1, core 4, gassum formation (see fig. 10a, 0.5 m). f: top of calcite-cemented shoreface sandstones with shell fragments forming the lst of sequence 3 (vi 1 third-order sequence). gassum-1 well, gassum formation (fig. 13a, 26.5 m). g: finegrained, swaley cross-stratified and wave-rippled sandstones of the tst of sequence 3. farsø-1, gassum formation (fig. 10a, 7.2 m). h: two sets of swaley cross-stratified, finegrained sandstones, the upper set containing abundant organic debris and shell fragments in the lower part. børglum-1, core 20, gassum formation, hst of fj 1 (fig. 5, 0.7 m; see also fig. 20). i–l: fine-grained heteroliths of the offshore to shoreface transition zone showing a variable degree of wave-reworking and bioturbation. gassum formation. m: weakly laminated to homogenous, offshore mudstones with thin-shelled bivalves marking mfs 3, sequence 3 of vi 1. farsø-1, gassum formation (fig. 10a, 13.2 m). 478 clay si sand 2 4 mm 5 m 4 3 2 1 0 forced regressive systems tract sp-log flyvbjerg-1 core 6 highstand systems tract core 7 clay si sand 2 4 mm marine flooding surface lagoonal flooding surface sp-log 4 m 0 3 2 1 core 8 lacustrine flooding surface lacustrine flooding surface clay si sand 2 4 mm sp-log 5 m 0 4 3 2 1 transgressive systems tract upper shoreface shoreface shoreface – offshore transition zone lagoon fluvial channel? lake swamp lake lagoon ✩ fig. 12. core logs of the gassum formation in the flyvbjerg-1 well. core 8 shows lagoonal and lacustrine deposits (tst) of the fourth-order sequence 7 of ga 1 (see fig. 19). core 7 shows fluvial sandstones overlain by lagoonal sandstones and marine heteroliths of the highly variable hst of sequence fj 1 (see fig. 20). core samples yielded ostracods, dinoflagellate cysts, spores and pollen indicating a hettangian age (michelsen 1975; poulsen 1992b). core 6 shows shoreface sandstones of the frst of sequence fj 2; core samples yielded spores and pollen of the sinemurian c. macroverrucosus zone (poulsen 1992b). for legend, see fig. 5. 479 clay si sand 2 4 mm transgressive systems tract lowstand systems tract transgressive systems tract vinding fm gassum fm rsme ✩ gassum-1 cores 67–75 highstand systems tract forced regressive systems tract hst lst hst gamma-ray offshore marine offshore marine offshore upper shoreface upper shoreface 32 m 42 12 1 30 29 27 25 20 15 11 10 0 mfs 3 tsme/ts 3 sb 3 mfs 2 rsme/sb 2 a this page and overleaf: fig. 13. core logs (a–c) of the vinding, gassum and fjerritslev formations in the gassum-1 well. the cores show the part of the hst of sequence vi 1 that is made up of the fourth-order sequences 1–4. vi 1 is overlain by sequence ga 1 (sb 5 – sb 9) consisting of the fourth-order sequences 5–8. the succeeding strata (above sb 9) are referred to sequence fj 1. the ammonite curviceras sp. of the upper planorbis zone occurs at 44 m. core depths are corrected to log depths by subtracting 20–25 ft. for legend, see fig. 5. 480 sb 6 sb 5 mfs 5 gassum-1 core 57–66 ts 5 lst transgressive systems tract 31 m 29 22 20 14 5 0 17 transgressive systems tract highstand systems tract gassum fm lowstand systems tract highstand systems tract frst gamma-ray offshore marine shoreface offshore marine offshore marine shoreface estuarine shoreface – offshore transition zone shoreface clay si sand 2 4 mm b 481 gamma-ray gassum-1 core 52–56 50 m 48 46 44 25 23 20 18 10 8 6 2 0 sb 9 mfs 8 ts 9 mfs 7 sb 7 mfs 6 tsme tsme/ts 8 rsme tsme sb 8 offshore shoreface offshore shoreface foreshore shoreface lake foreshore– shoreface offshore offshore fjerritslev fm gassum fm transgressive systems tract lowstand systems tract highstand systems tract transgressive systems tract lowstand systems tract highstand systems tract transgressive systems tract lowstand systems tract forced regressive systems tract clay si sand 2 4 mm c 482 gamma-ray clay si sand 2 4 mm estuarine fill stenlille-1 cores 9, 10 25 m 20 15 10 5 0 lagoon lagoonal flooding surface estuarine shoal tidal creeks erosional channel base channel bar/margin tidal channel lowstand systems tract transgressive systems tract sb 5 ts 5 fig. 14. core log of the gassum formation in the stenlille-1 well. the cores show the lst of the fourth-order sequence 5 of the ga 1 sequence consisting of an estuarine channel-fill, which is overlain by lagoonal deposits resting on a bay-line transgressive surface interpreted as ts 5. (log measured by l. hamberg; modified from l.h. nielsen, l. hamberg and e.b. koppelhus in: nielsen 1995). for legend, see fig. 5. sandstones with carbonaceous debris and common mudstone drapes on foresets, occasionally forming couplets (nielsen et al. 1989; hamberg 1994; l.h. nielsen, l. hamberg and e.b. koppelhus in: nielsen 1995). the foresets are commonly tangential and pass into cross-laminated bottomsets. ripples occur locally in the lower part of the foresets with foreset laminae dipping in the opposite direction to the larger foresets. interbedded with the cross-beds are flaser to wavy bedded heterolithic beds, 2–15 cm thick, with erosional bases and sharp boundaries between ripple-laminated, non-graded sandstone layers and mudstone drapes. upwards in the succession, the cross-beds are overlain by fining-upwards beds, 10–25 cm thick, with small, cross-bedded sets that progressively decrease in thickness to ripple-lamination capped by mudstone or flaser bedding. higher in the section, erosionally based beds show inclined heteroliths of thinly interlaminated, carbonaceous sandstone and mudstone. the inclined heteroliths are overlain by weakly fining-upwards units with erosional bases showing ripple-lamination, parallel lamination and flaser bedding. these units and the inclined heteroliths are burrowed and contain abundant comminuted organic debris, and fragments of leaves, wood and probable equisetites sp. stems. the overlying, slightly heterolithic and burrowed sandstones show mainly parallel lamination and muddraped wave-ripple lamination. the succession ends in a burrowed, muddy layer with ophiomorpha isp., rootlets, stems and desiccation cracks. the mixed marine and freshwater palynomorph assemblages, burrowing, abundant plant material, rootlets and coal seams indicate a brackish-marine environment with fluvial influx. the regular occurrence of foreset drapes reflects tidal influence during deposition, and couplets indicate subtidal conditions (visser 1980; banerjee 1989). the cross-beds were formed by migrating transverse bedforms during dominant tidal currents, while the ripple-laminated sand enclosed by mud couplets reflects deposition by the oppositely directed, subordinate and weaker current causing ripple migration up the slipface of the larger bedforms (boersma & terwindt 1981; terwindt 1981). deposition is thus interpreted as having taken place in channels with relatively strong tidal currents. the sharp boundaries between the sandstones and the mudstone partings in the heteroliths indicate a settling time-lag of mud relative to sand deposition, which is characteristic of tidal heteroliths (van straaten & kuenen 1957; little-gadow & reineck 1974), and the heteroliths are interpreted as tidal flat deposits. the increase upwards in the proportion of mud reflects progradation of the tidal flats. the bioturbated, inclined heterolithic bedding is interpreted to have formed by lateral accretion on point bars in intertidal creeks (allen 1965; barwis 1978; thomas et al. 1987; smith 1988). the uppermost heteroliths and mudstones with rootlets in the stenlille-1 core reflect upwards shallowing and eventual subaerial exposure. log correlation between closely spaced wells at stenlille indicate the presence of a deeply incised valley, and the stenlille-1 core is interpreted to represent the fill of a deep estuarine channel (hamberg 1994; l.h. nielsen, l. hamberg and e.b. koppelhus in: nielsen 1995). lagoonal facies association this association occurs commonly in the gassum formation and was cored in the flyvbjerg-1, frederikshavn-2, stenlille-1, -2, -6, thisted-3, vedsted-1 and års-1 wells (figs 7, 8, 9, 12, 14, 15). it is locally present in the uppermost part of the fjerritslev formation on the skagerrak–kattegat platform, where it was cored in the skagen-2 well. it is relatively uncommon in the haldager sand formation, and was only cored in the frederikshavn-1 well. the association consists of muddy and sandy heteroliths and carbonaceous claystones and sandstones, typically forming successions 1.5–11 m thick, capped by coal seams, carbonaceous claystones and heteroliths with rootlets or a ravinement surface with shoreface sandstones. good examples are evident in cores from the gassum formation at stenlille (fig. 14; hamberg 1994; l.h. nielsen, l. hamberg and e.b. koppelhus in: nielsen 1995). estuarine coaly beds are erosionally overlain by wave-rippled, carbonaceous sandstones and heteroliths of mudstone and well-sorted, fine-grained sandstone and siltstone showing parallel to incipient wave-ripple lamination or hummocky cross-stratification. abundant sand-filled synaeresis cracks occur in the mudstones. the sandstones thin and fine upwards, passing up into a laminated heterolith of mudstone and siltstone. the trace fossils teichichnus isp. and planolites isp. occur commonly. other examples show laminated and wave-rippled siltstones and heteroliths interbedded with erosionally based cross-bedded sandstones and inclined heterolithic strata (figs 7a, 9b). drapes of organic debris and small ripple-foreset laminae occur locally within the foresets. the successions fine upwards from fineto mediumgrained sandstones to highly carbonaceous heteroliths or muddy siltstones, commonly homogenised by dense rootlets. siderite-cemented layers, 1–5 cm thick, commonly occur in the mudstone beds. the inclined het483 eroliths dip 2–10° and are composed of thinly and rhythmically interbedded carbonaceous sandstones and mudstones with large fragments of plants and branches. a sapropelitic coal bed, 5 cm thick, with a very high content of alginite, sporinite and bituminite occurs in the års-1 core (fig. 9b; thomsen 1984). from the uppermost part of the fjerritslev formation as defined by well logs in skagen-2 (michelsen 1978), cores with poor recovery show units, 2–4 m thick of greyish, fine-grained micaceous sandstones and coarse siltstones, dark grey to almost black, carbonaceous and poorly laminated mudstones, and coaly beds. thick, long rootlets are very common; an upwards change from very small to large rootlets occurs in one unit. cores with poor recovery from the haldager sand formation in the frederikshavn-1 well show 10–12 m of almost homogeneous dark claystones with few silt-streaks, muddy and sandy heteroliths with laminated and muddraped cross-laminated sand layers, and a few thin coarse-grained, pebbly layers. pyrite occurs as large nodules and as cement in sandstones. plant fragments, large coal fragments or coal seams, possible rootlets and weak bioturbation are also present. the wave-rippled sandstone at the base of the lagoonal succession in figure 14 is interpreted as a transgressive lagoonal shoreline deposit, and the overall fining-upwards trend of the wave-reworked sandstones and heteroliths reflects deposition in increasingly deeper lagoonal water. facies successions capped by coal beds or rooted sediments are interpreted as the fill 484 clay si sand 2 4 mm sp-log sp-log frederikshavn-2 cores 3, 4 core 6 transgressive systems tract transgressive systems tract lowstand systems tract transgressive systems tract sb 15 ts 13 ts 15 clay si sand 2 4 mm lower shoreface shoreface offshore fluvial lagoon swamp lagoon swamp la. fs la. fs/ts 9 offshore shoreface lower shoreface – offshore transition zone 3 2 1 0 4 m 5 8 m 4 3 2 1 4 m 3 2 1 0 clay si sand 2 4 mm core 7 transgressive systems tract lowstand systems tractsp-log fig. 15. core logs of the gassum (core 7, lower core 6) and fjerritslev formations (upper core 6, cores 4, 3) in the frederikshavn-2 well. core 7 shows fluvial deposits of the lst overlain by lagoonal deposits of the tst, sequence fj 1. core 6 shows sandstones and heteroliths of the shoreface – offshore transition zone (gassum fm) and offshore mudstones (fjerritslev fm) belonging to the tst of sequence fj 3. the mudstones have yielded d. priscum dinoflagellate cysts (dybkjær 1991). core 3 illustrates a thin calcite-cemented sandstone bed that overlies, and is succeeded by, offshore mudstones. ostracods of the g. apostolescui – k. (k.) foveolata subzone occur in the transgressive, offshore mudstones of sequence fj 5 (between ts 13 and sb 15) indicating an early pliensbachian age, while ostracods of the upper part of the o. adenticulata – n. (n.) simplex zone occur in the transgressive mudstones of sequence fj 7 above ts 15 indicating a late late pliensbachian age (michelsen 1975). a hiatus corresponding to at least the margaritatus zone is thus indicated at sb 15. la. fs, lagoonal flooding surface. for legend, see fig. 5. of lagoons that gradually developed into coastal marshes (e.g. fig. 15, core 7), while successions overlain by shoreface sandstones reflect rising sea level and marine flooding of the coastal area. the fining-upwards units of erosionally-based cross-bedded sandstones overlain by carbonaceous heteroliths or siltstones with rootlets are interpreted as the fills of small channels, and the inclined, stratified heteroliths were formed by point bar deposition (thomas et al. 1987). deposition probably took place in tidally influenced creeks that formed part of a tidal drainage network near a vegetated lagoonal margin. a brackish lagoon is similarly indicated by the weakly bioturbated sandstones and mudstones with abundant rootlets and well-preserved plant material in the flyvbjerg-1 and vedsted-1 cores (figs 8, 12). the composition of the coal bed in års-1 indicates deposition of organic-rich mud in an anoxic brackish to freshwater environment. the successions cored from the uppermost part of the fjerritslev formation and the lowermost part of the haldager sand formation in the skagen-2 well, and from the middle part of the haldager sand formation in the frederikshavn-1 well are interpreted as lagoon and swamp deposits by analogy to well-cored examples from contemporaneous deposits on bornholm (nielsen 1995). the interpretation is supported by a significant pyrite content in the coal seam from the fjerritslev formation (petersen et al. 2003, this volume). the delicate rootlets may represent small, salt-tolerant pioneer plants or sea-meadows that vegetated muddy lagoonal flats, while the larger rootlets close to the top of the succession may represent larger land plants that became dominant when the lagoon silted up and became a marsh. shoreface and foreshore facies association this association is very common in the gassum formation and has been cored in many wells (børglum-1, farsø-1, flyvbjerg-1, frederikshavn-2, gassum-1, horsens-1, stenlille, thisted-3, ullerslev-1, vedsted-1 and års-1 wells; figs 5–7, 10, 12, 13, 15). the association occurs locally in the fjerritslev formation especially from the sorgenfrei–tornquist zone and the skagerrak–kattegat platform. the facies association consists primarily of hummocky cross-stratified and wave-rippled, light olivegrey to grey, locally weakly glauconitic sandstones and coarse-grained siltstones with coal fragments and scattered pyrite nodules in places (fig. 11c–k). in the gassum formation, the association is 4–30 m thick and typically sharply overlies offshore mudstones (l.h. nielsen, l. hamberg and e.b. koppelhus in: nielsen 1995). thick sandstone successions are common close to the southern and eastern basin margins, and in the basin centre in the lower part of the gassum formation. the sandstones are erosionally based and either show blocky log motifs or comprise several units, 2–5 m thick, separated by thinner and slightly finer grained sandy beds or, in places, heteroliths. the heteroliths may be intensely burrowed and teichichnus isp. is present. the gassum-1 core from a blocky succession shows faint, parallel lamination, low-angle cross-bedding and structureless intervals (fig. 13a). wave-ripple crosslamination and trough cross-bedding may be outlined by drapes of organic debris. the sandstones are generally capped by a sharp transgressive surface of marine erosion overlain by offshore mudstones, or locally by a calcite-cemented sandstone with small pebbles and shells or up to 0.5 m of fining-upwards sandstones and heteroliths (fig. 11d–f). in the upper part of the gassum formation, the sandstones are thinner and typically form two or three units separated by muddy, burrowed, wave-rippled and wavy bedded heteroliths (hamberg & nielsen 2000). horizontal burrows, large vertical burrows and skolithos isp. occur locally in the sandstones. the lower sandstones are typically fine-grained and hummocky crossstratified with a sharp, erosional base mantled by coarser grains, rip-up clasts, sideritic clasts and carbonaceous debris. the upper sandstones are fineto mediumgrained, commonly with a weak coarsening-upwards tendency, and show swaley cross-stratification, trough and planar cross-bedding, low angle cross-bedding, parallel lamination and wave-ripple lamination. rootlets are locally preserved (e.g. stenlille-2; hamberg & nielsen 2000). sharp-based shoreface sandstones also occur in the fjerritslev formation as shown by cores and well logs in the stenlille wells (figs 16, 17). cuttings samples and well logs from the sorgenfrei–tornquist zone and the skagerrak–kattegat platform indicate a succession of sandstones or siltstones interbedded with marine mudstones. core pieces from the upper part of the f-ii member of the fjerritslev formation from the skagerrak–kattegat platform (frederikshavn-1, -2) show an olive-grey, greenish or weakly yellowish, micaceous and carbonaceous sandstone which is very fineto finegrained, muddy, occasionally heterolithic, and almost completely burrowed. weak parallel to wavy lamination and cross-lamination are locally present. shell frag485 486 box 11 1530.20–1531.20 m box 12 1531.20–1532.20 m box 13 1532.20–1533.20 m box 14 1533.20–1534.20 m box 15 1534.20–1535.20 m 0 10 20 30 40 50 60 70 80 90 cm 100 ts 9 tsme sb 9 rsme 487 ments occur abundantly and thin sideriteand calcitecemented beds are common. core pieces from the upper part of the f-iv member (frederikshavn-1, skagen-2) show an olive grey, micaceous and carbonaceous, poorly cemented, coarse-grained siltstone grading to very fine-grained sandstone with thin mudstone laminae. the trace and body fossils, glauconite and pyrite indicate marine conditions. the sharp, basal erosion surface of the sandstones is interpreted as the result of prolonged wave scouring and winnowing in the shoreface during sea-level fall causing bypass and seawards transport of fine-grained material, and formation of a thin lag of coarse sand, small pebbles, claystone chips and plant fragments. the overlying amalgamated hummocky cross-stratified sandstone beds were deposited in the lower shoreface during the waning phase of storms, which eroded away any fair-weather, fine-grained deposits (dott & bourgeois 1982). the interbedded heteroliths are interpreted as having been deposited in the offshore–shoreface transition zone. the trough crossbedding to swaley cross-stratification and the association of wave-rippled and cross-bedded beds suggest constant reworking by waves and strong currents in the upper shoreface and foreshore (davidson-arnott & greenwood 1976; hunter et al. 1979; leckie & walker 1982; nielsen et al. 1988). the planar cross-beds are interpreted as swash-bars, and the parallel lamination and low-angle cross-bedding represent beach lamination (thompson 1937; clifton 1969; clifton et al. 1971). the vertical facies successions represent shoreline progradation with minor intermittent transgressions represented by the intercalated heteroliths or fine-grained sandstone beds. heteroliths are sometimes absent, suggesting that the entire sandstone unit was deposited in the upper shoreface to foreshore (e.g. gassum-1; fig. 13a). with the exception of the scattered rootlet horizons, no evidence of subaerial exposure is seen and the erosion surface at the base of the overlying marine mudstones is interpreted as a ravinement surface. the sandstones of the f-ii and f-iv members of the fjerritslev formation were deposited in the offshore–shoreface transition zone as deduced from their fine-grained, muddy and heterolithic nature. the pervasive bioturbation and abundant bivalves in the f-ii member indicate a well-oxygenated sea-bottom with infaunal deposit-feeders and suspension-feeders similar to those described from other parts of the fjerritslev formation (pedersen 1986). the lack of shells and burrows in the f-iv sandstones may relate to the generally unfossiliferous nature of the f-iv member, a feature that has been interpreted to reflect anoxia (michelsen 1975, 1989b). ts 11 tst lst fj 3 sb 11 offshore mudstone shoreface sandstone fj 2hst depositional environment bounding surfaces sb sequence boundary ts transgressive surface correlation line stenlille-1 gr stenlille-2 gr stenlille-5 gr stenlille-4 gr fig. 17. log panel of the fjerritslev formation in the stenlille structure, showing marine shoreface sandstones of the lst of sequence fj 3, 10–15 m thick, sharply overlying marine offshore mudstones of the hst of fj 2. sb 11 is developed as a regressive surface of marine erosion. the sandstones are capped by ts 11, developed as a transgressive surface of marine erosion, overlain by marine mudstones of the tst. the wells are placed in a hypothetical cross-section (c. 1 km long) perpendicular to the coast with land to the right and sea to the left. the sandstones wedge out seawards and the sandwiched mudstones thicken. facing page: fig. 16. marine lower shoreface to offshore transition zone heteroliths of the hst of sequence ga 1 (boxes 13 lower part, 14, 15) erosionally overlain by shoreface sandstones (boxes 12, 13 upper part) of the lst of sequence fj 1 bounded below by sb 9, which is developed as a regressive surface of marine erosion (rsme). the lst is sharply topped by ts 9, a transgressive surface of marine erosion (tsme), overlain by transgressive heteroliths and offshore mudstones (boxes 11, 12) of the tst of sequence fj 1. f-1a member, stenlille-5, core 3; base lower right, top upper left. recent sandy shoreface profiles commonly show a break in slope from the shoreface sand to offshore mud at water depths of 5–15 m (everts 1987; penland et al. 1988; walker & plint 1992; nummedal et al. 1993); shoreface progradation during a sea-level stillstand would thus produce 5–15 m of sand. the thick units in the gassum formation that show several sandstone beds with finer-grained intercalations are therefore interpreted as amalgamated progradational shoreface sandstones stacked in an aggradational manner, reflecting rising sea level during deposition. detailed correlation of closely-spaced cored sections in the stenlille area indicates that the sandstones with sandwiched mudstones and heteroliths are composed of shingles formed during forced regressions punctuated by minor transgressions that caused deposition of the fine-grained deposits (hamberg & nielsen 2000). marine offshore facies association the association is common in the gassum formation where it has been cored in the farsø-1, gassum-1, stenlille, ullerslev-1 and års-1 wells (figs 5, 8–10, 13, 15). it dominates the fjerritslev formation and has been cored in many wells (fig. 16; michelsen 1975, 1978, 1989a, b; pedersen 1985, 1986). cores and log patterns indicate that it occurs in the haldager sand formation from the sorgenfrei–tornquist zone. the association is dominated by mudstones but also includes siltstones, heterolithic siltstones and very finegrained, heterolithic sandstones. the mudstones are black to dark grey, carbonaceous, with graded or uniform claystone laminae. scattered silt streaks and lenses are common. small planolites isp. and chondrites isp. burrows are locally present. the mudstones grade into silt-streaked mudstones with dark grey to olive grey, erosionally based, graded siltstone laminae. incipient ripples are also present; they are 5–10 mm thick, composed of coarse siltstone or very fine-grained sandstone with internal flat cross-lamination, and, in places, loading. indistinct bioturbation is observed. the heterolithic siltstones consist of black mudstones and olive grey siltstones grading to very fine-grained sandstones. the siltstone laminae are up to 3 cm thick, graded and have sharp, erosional bases (fig. 11i, j). steep-sided scours are observed in places; they are typically 1–4 cm deep and filled with silt or very fine-grained sand. wave ripple cross-lamination, with chevrons, offshoots and draping foresets, is common. burrows are common and the heterolithic siltstones are sometimes completely churned (fig. 11k, l). the trace fossils zoophycos isp., thalassinoides isp., teichichnus isp. and helminthopsis isp. occur sporadically. very fine-grained, heterolithic sandstones, constructed of sand laminae up to 2 cm thick and thinner mudstone laminae, show wavy-bedding that is almost obliterated by bioturbation. well-sorted, coarse-grained siltstones and very fine-grained sandstones show small-scale hummocky cross-stratified sets, 5–15 cm thick, separated by mudstones. the mudstones typically form relatively uniform successions that may be several tens of metres thick. in places, they form 1–5 m thick units that become siltier upwards, beginning with black mudstones and terminating with hummocky cross-stratified siltstones, which in turn are abruptly overlain by black mudstones. coarsening-upwards successions of silt-streaked mudstones overlain by wavy-bedded, heterolithic sandstones also occur (fig. 5, core 17). the sandstones are capped by a graded, fine-grained sandstone layer containing quartz pebbles up to 0.5 cm in diameter and overlain by a dark grey–black mudstone, which may be carbonate-cemented. coarsening-upwards successions, 3–5 m thick, overlain by fining-upwards successions, 5–15 m thick, have been described from the fjerritslev formation (pedersen 1985). in the gassum formation, the facies association contains dinoflagellate cysts, acritarchs, foraminifers, rare marine bivalves, and fragments of belemnites and ammonites. in the fjerritslev formation, it contains a rich fauna of ostracods, bivalves, gastropods, foraminifers, ammonites and other marine fossils in addition to dinoflagellate cysts (nørvang 1957; sorgenfrei & buch 1964; michelsen 1975; pedersen 1986; dybkjær 1988, 1991; poulsen 1996). in the haldager sand formation, dinoflagellate cysts, tasmanites (prasinophycean algae), acritarchs and foraminifers are found (forbes et al. 1985; ravn-sørensen 1989; poulsen 1992a). the body and trace fossils and palynomorphs clearly indicate a marine environment. the nature of the mudstones testifies to deposition primarily in calm water from suspension clouds. the graded bedding, scoured surfaces and incipient ripples indicate rapid deposition from the suspension load of storm-generated currents (pedersen 1985). isolated coarse silt to fine-grained sand ripples, wave ripples and hummocky cross-stratification reflect an increasing amount of wave reworking during storm events (de raaf et al. 1977). the scour-fills are interpreted as small gutter-casts reflecting scouring and deposition. the environment was offshore marine below average fair-weather wave base ranging from below storm wave base (the mudstone 488 end-member) to the transition zone influenced by storm waves (the hummocky cross-stratified end-member). pedersen (1986) recognised a normal shale facies fauna in the hettangian and lower sinemurian, and a restricted shale facies fauna in the upper sinemurian – pliensbachian on the basis of fossil content and bivalve ecology in cores from the gassum-1 well. the black, carbonaceous and clay-dominated mudstones characterised by the highest gamma-ray values and maximum (leftward) deflection of the sp-log are interpreted to represent the lowest accumulation rate during maximum flooding. the thick, uniform successions in the fjerritslev formation are interpreted to reflect relatively rapid vertical aggradation of mud blankets deposited from the suspension loads of storm-generated currents. the more distinctly coarsening-upwards units encased in black mudstones are interpreted as distal parasequences formed by coastal progradation terminated by flooding. the thin, fine-grained pebble layer is interpreted as the distal correlative of a ravinement surface formed during the transgression of the coastal areas, and the carbonate cement probably reflects early diagenesis associated with condensation caused by subsequent flooding. the association corresponds to the ‘outer shelf facies association’ and the most fine-grained part of the ‘inner shelf facies association’ described from the fjerritslev formation by pedersen (1985). depositional sequences the upper triassic – jurassic of the danish basin is composed of depositional sequences falling in the time range of the secondto third-order sequences of mitchum & van wagoner (1991). for convenience, they are here termed third-order sequences without implying correlation to the cycle chart of haq et al. (1988). representative well-log cross-sections of the sequences are shown in figures 19–22, 24 and 25. the third-order sequences are numbered sequentially within each formation, from vi 1 (vinding formation) at the base of the succession to fr 3 (frederikshavn formation) at the top; the uppermost jurassic to lower cretaceous bø 1, fr 1, fr 2 and fr 3 sequences are not described in detail here. the vi 1 and ga 1 third-order sequences are further composed of fourth-order sequences; these are described and shown on figures 20 and 21 but are not named. sequence boundaries (sb), maximum flooding surfaces (mfs) and transgressive surfaces (ts) are numbered from 1 (lower) to 23 (upper) irrespective of their hierarchy, although only third-order surfaces are shown on figure 25. the facies associations described above form the basis for interpretations of well logs and the subdivision into depositional units. sequence boundaries, maximum flooding surfaces and transgressive surfaces are identified following van wagoner et al. (1990). the sequence stratigraphic interpretation of deposits formed during sea-level fall and the position of the sequence boundary is intensely debated (plint 1988; posamentier et al. 1988, 1992; hunt & tucker 1992, 1995; ainsworth & pattison 1994; helland-hansen & gjelberg 1994; kolla et al. 1995; hamberg & nielsen 2000; plint & nummedal 2000; posamentier & morris 2000). in this study, the shelf deposits formed during falling sea level are included in the forced regressive systems tract as defined by hunt & tucker (1995). fluvial incision is interpreted to have occurred during sea-level fall, although deposition may also occur during the fall (van wagoner 1995). the fluvial and estuarine valley-fill deposits overlie incised surfaces and are bounded above by regional marine flooding surfaces with marine mudstones. the valley-fills are thus interpreted to have formed during sea-level lowstand and early phases of sea-level rise, and are therefore included in the (late) lowstand systems tract (van wagoner et al. 1990), rather than the (early) transgressive systems tract as preferred by others (allen & posamentier 1994). the dating of the sequences is primarily based on ostracod, dinoflagellate and miospore zones (fig. 18). the standard ammonite zones are used as chronostratigraphic units (callomon 1984) and are referred to by the species name alone – the turneri zone, for example. the resolution of seismic data from the danish basin is in general too poor to confirm the well-log correlations. the reflectors are relatively parallel except for areas influenced by growth of salt structures, and attributes such as downlap, onlap and offlap are difficult to recognise. it is furthermore difficult to correlate reflectors across major faults and between onshore and offshore surveys. some regional seismic reflectors may, however, be identified to support the log correlations. the ‘top oddesund formation’ reflector seems to correspond to a significant lower norian transgressive surface at the base of marine claystones of the vinding formation. the ‘top triassic’ reflector, which forms a very distinct seismic marker traceable over most of the basin (baartman & christensen 1975; japsen & langtofte 1991), seems to correspond to an upper rhaetian maximum flooding surface (l.h. nielsen, l. hamberg and e.b. koppelhus in: nielsen 1995). the seismic marker corresponding to the top of the f-ii member is a 489 490 poulsen 1996 c. betzi – c. crassireticulata r. rhaetica n . g ra ci lis l. variabile d. priscum dinoflagellate cyst zones marshi reticulatus quinquepunctatus macer zigzag parkinsoni garantiana subfurcatum humphriesianum sauzei laeviuscula discites concavum murchisonae opalinum aalensis pseudoradiosa dispansum thouarence variabilis bifrons falciferum tenuicostatum spinatum margaritatus davoei ibex jamesoni raricostatum oxynotum obtusum turneri semicostatum bucklandi angulata liasicus planorbis chronostratigraphy barren poor records ostracod zones ricciisporites–polypodiisporites rhaetipollis–limbosporites ricciisporites–conbaculatisporites corollina–enzonalasporites pinuspollenites– trachysporites cerebropollenites macroverrucosus spheripollenites– leptolepidites perinopollenites elatoides miospore zonespoulsen & riding 2003 r. rhaetica p. nasuta n. gracilis c. crispum l. spinosa m. reticulatum l. variabile d. priscum n. senex m. semitabulatum p. nasuta not zoned m. semitabulatum b a b b a a l. spinosa m. semitabulatum not zoned o. danica o. aspinata o. adenticulata – n. (n.) simplex ‘lower subzone’ ‘middle subzone’ p. reticulata ‘upper subzone’ g. apostolescui – k. (k.) foveolata ‘upper subzone’ ‘lower subzone’ rhaetian norian (part) bath. (part) bajocian aalenian toarcian pliensbachian sinemurian hettangian stages zone subzone significant upper pliensbachian transgressive surface. the ‘base middle jurassic unconformity’ is a deep truncation surface on the ringkøbing–fyn high and in the south-western part of the basin caused by uplift (michelsen 1978; michelsen & andersen 1981; ziegler 1982; koch 1983; eugeno-s working group 1988). less section is missing below the unconformity in the basin centre, and it continues into a seismic conformity in the sorgenfrei–tornquist zone. sequence vi 1 (lower norian – lower rhaetian) this sequence comprises the vinding formation, the upper part of the skagerrak formation and the lower part of the gassum formation. it spans the upper norian and rhaetian miospore assemblage zones i–iii of bertelsen (1978; zone l was dated to the rhaetian by bertelsen, but is of late norian age according to pedersen & lund 1980). furthermore, miospore zones ranging from the corollina–enzonalasporites zone to the lower part of the rhaetipollis–limbosporites zone and the dinoflagellate rhaetogonyaulax rhaetica zone have been identified in some wells (fig. 18; l.h. nielsen, l. hamberg and e.b. koppelhus in: nielsen 1995). these data indicate an early norian – early rhaetian age. this is supported by the presence of the foraminifer ammodiscus sp., the ostracod emphazia sp. and the bivalve rhaetavicula contorta (fazekas 1948; fredbold 1948; nørvang 1948; sorgenfrei & buch 1964; christensen 1972). base of sequence and transgressive systems tract the western, central and southern parts of the danish basin were flooded during the early norian, and continental–sabkha deposits of the oddesund formation were overlain by restricted and shallow marine deposits of the vinding formation (christensen 1972; bertelsen 1978, 1980). the boundary between the two formations is interpreted as a transgressive surface ts 1, and is tentatively adopted as the base of the vi 1 sequence (fig. 19, facing page 498). the transgression culminated in the late norian (bertelsen 1978), as marked by the maximum flooding surface mfs 1 that is readily recognised in the vinding formation on sp and gamma-ray logs from most wells. the mfs 1 is characterised by an olive-grey, very finely laminated claystone in cores (e.g. ullerslev-1). mfs 1 is tentatively traced into thick lacustrine claystones occurring in the upper part of the skagerrak formation along the north-eastern and eastern basin margins and in the western part of the basin (f-1, inez-1, k-1; fig. 1). the lacustrine claystones are assumed to be contemporaneous with the marine claystones in the vinding formation based on biostratigraphic evidence (bertelsen 1980); the sporadic occurrence of acritarchs (bertelsen 1978) suggests that the lakes were slightly brackish and probably formed as a result of the general rise in base-level caused by rising sea level. the transgressive systems tract from ts 1 to mfs 1 consists of fossiliferous limestones, oolites and marine claystones, up to 60 m thick in the basinal areas. south of the ringkøbing–fyn high, in the north german basin, the vinding formation is similarly developed indicating that the high was submerged at this time, precluding the supply of coarse sediment. highstand systems tract the highstand systems tract bracketed by the thirdorder surfaces mfs 1 and sb 5, consists of a weakly coarsening-upwards succession of marine mudstones, 10–15 m thick, overlain by alternating marine mudstones, shoreface sandstones and estuarine deposits amounting to a total thickness of the hst of up to 125 m (fig. 19). the shoreface sandstones are typically sharp-based and occur together with the estuarine deposits at three distinct levels separated by marine mudstones. the three successive shoreface sandstones show a stepwise more basinwards distribution, and define three forestepping fourth-order sequences bounded by sb 2, 3 and 4 (l.h. nielsen, l. hamberg and e.b. koppelhus in: nielsen 1995; hamberg & nielsen 2000). the sandstones overlie wave-scoured surfaces and were mainly deposited during falling sea level and constitute widespread, forced regressive systems tracts. a cored example is shown in figure 13a from gassum-1. in some places, the shoreface sandstones are 491 facing page: fig. 18. biozonation of the danish basin and the fennoscandian border zone based on data from michelsen (1975, 1989a), dybkjær (1988, 1991), poulsen (1992a, 1996), l.h. nielsen, l. hamberg and e.b. koppelhus (in: nielsen 1995), koppelhus & batten (1996) and poulsen & riding (2003, this volume). up to 30 m thick and show a subtle forestepping to aggrading pattern overlain by an aggrading to backstepping pattern reflecting weak overall shoreface progradation during late highstand to falling sea level (highstand to forced regressive systems tract) followed by overall shoreline retreat during the subsequent rise (lowstand to early transgressive systems tract; l.h. nielsen, l. hamberg and e.b. koppelhus in: nielsen 1995). the shoreface sandstones were incised in places due to further sea-level fall; during subsequent sealevel rise, estuarine sediments were deposited in the incised valleys. the vi 1 sequence seems to be absent on the skagerrak–kattegat platform. in the fjerritslev trough, which is down-faulted relative to the platform, the sequence consists of alternating units of pebbly fluvial sandstones and lacustrine clays of the upper skagerrak formation and lower gassum formation (fig. 19). sequence ga 1 (lower rhaetian – lower hettangian) this sequence is primarily composed of the upper part of the gassum formation. the sequence also includes the lowermost part of the fjerritslev formation in the basin centre, and 10–20 m of greyish claystones of the uppermost skagerrak formation on the skagerrak– kattegat platform which probably belong to the miospore rhaetipollis–limbosporites zone (e.b. koppelhus, personal communication 1994). with reference to the dinoflagellates, the sequence spans the upper part of the rhaetogonyaulax rhaetica zone and the lowermost part of dapcodinium priscum zone; according to the spores and pollen, it spans the upper part of the rhaetipollis–limbosporites zone, the ricciisporites– polypodiisporites zone and the lowermost part of the pinuspollenites–trachysporites zone. base of sequence and lowstand to transgressive systems tract the lower boundary of the third-order sequence is defined by a basinwide, lower rhaetian fluvial incision surface, sb 5, present in the lower gassum formation (fig. 19). the third-order sb 5 marks the largest basinwards shift in facies and the most pronounced fluvial incision in the upper norian – rhaetian succession, and an extensive late rhaetian coastal onlap onto the surface occurs towards the north-east (l.h. nielsen, l. hamberg and e.b. koppelhus in: nielsen 1995). the sequence boundary cuts into marine offshore mudstones and shoreface sandstones of the lower gassum formation. in the deep central part of the basin, the sequence boundary occurs at the base of shoreface sandstones (e.g. voldum-1) or at the base of estuarine sandstones (e.g gassum-1; fig. 13b). close to the southwestern basin margin, along the ringkøbing–fyn high (horsens-1, ullerslev-1, vemb-1) and the north-eastern margin at the børglum fault (flyvbjerg-1), the sequence boundary is overlain by fluvial sandstones belonging to the lowstand systems tract. in the himmerland graben (farsø-1, års-1), the sorgenfrei–tornquist zone (terne-1), at stenlille (stenlille-13) and at thisted (thisted-1, -3, -4), the sequence boundary shows deep incision and it is overlain by fluvial–estuarine deposits up to 45 m thick, showing a variable development. cores from the incised valley-fills show fluvial–estuarine and lagoonal sandstones and mudstones (figs 7a, b, 8, 9a, 14). the valley-fills commonly have a basal unit of fluvially dominated sandstones, a middle unit of lagoonal mudstones and an upper unit of shoreface sandstones reflecting vertical stacking of an estuarine valley system caused by progressive drowning (allen & posamentier 1994; dalrymple et al. 1994). in the stenlille area, the thick incised valley-fill is capped by a heterolithic sandstone with roots overlain by transgressive lagoonal deposits (fig. 14). in some wells, the sequence boundary occurs on top of shoreface sandstones belonging to the underlying forced regressive systems tract (e.g. rønde-1, voldum-1) and it coincides with a transgressive surface of marine erosion indicating very limited fluvial incision. the transgressive systems tract overlying the transgressive surface ts 5 typically consists of 20–40 m of offshore mudstones and shoreface sandstones. the systems tract is relatively uniformly developed over large parts of the basin and shows an overall backstepping pattern that culminates with a widespread marine claystone containing the third-order mfs 7. two widely distributed shoreface sandstones occur within the transgressive systems tract; they both overlie a sharp, wavescoured base and are erosionally overlain by transgressive marine mudstones. in places, cores show roots preserved in shoreface sandstones below the ravinement surface indicating subaerial exposure during the lowest sea-level stand (hamberg & nielsen 2000). the sandstones are interpreted as forced regressive systems tracts reflecting two fourth-order sea-level falls superimposed on the general third-order rise that culminated at mfs 7 (sb 6, sb 7; fig. 19). the sandstones show a back492 stepping pattern and testify to a progressive, stepwise landwards shift of the lowstand shorelines. maximum flooding surface, mfs 7 and highstand systems tract the condensed section containing the mfs 7 is easily recognised on well logs, and is traced throughout the basin from the easternmost part (hans-1, lavø-1) to the westernmost (f-1, inez-1, k-1), from the skagerrak– kattegat platform (frederikshavn-2, sæby-1) to well sections south of the ringkøbing–fyn high (e.g. rødby-1), indicating that the entire basin including the platform was flooded. in cores from the gassum-1 well, the ammonite curviceras sp. (determined by g. bloos, s. elmi, s. franiatte and r. mouterde in: poulsen 1996), indicating the johnstoni subzone (the upper subzone of the lower jurassic planorbis zone), occurs c. 30 m above mfs 7 (fig. 13c). in rødby-1, where mfs 7 occurs in marine claystones and heteroliths marked by a distinct rightwards deflection of the sp-log, the triassic– jurassic boundary as determined by the spore–pollen ricciisporites–polypodiisporites and pinuspollenites–trachysporites zones occurs 4.3 m above mfs 7, and lower jurassic ammonites (psiloceras planorbis, determined by j.h. callomon and d.t. donovan in: poulsen 1996) and the base of the ostracod ogmoconchella aspinata zone are found 13 m and 2.4 m above mfs 7, respectively (bertelsen & michelsen 1970; michelsen 1973, 1975; lund 1977; poulsen 1992a, 1996). the lowermost hettangian ostracods and the rhaetian–hettangian boundary based on spores and pollen occur just above mfs 7 in several wells, and the boundary between the rhaetian r. rhaetica and the uppermost rhaetian – lowermost sinemurian d. priscum zones (dinoflagellates) seems to coincide with mfs 7 (michelsen 1975; l.h. nielsen, l. hamberg and e.b. koppelhus in: nielsen 1995; poulsen 1996). mfs 7 is thus well dated to the latest rhaetian and represents the maximum extent of the sea in rhaetian time. mfs 7 is overlain by an overall forestepping succession of marine mudstones and sandstones constituting a highstand systems tract, up to 35 m thick. cores from the stenlille area show offshore silt-streaked mudstones that coarsen upwards to lower shoreface heterolithic siltstones and sandstones. a higher-order sequence is recognised within the third-order highstand systems tract based on a widespread, thin shoreface sandstone that passes into fluvial sandstones overlying sb 8 in the north-eastern part of the basin (fig. 19). sequence fj 1 (lower–middle hettangian) sequence fj 1 consists of marine heteroliths of the uppermost gassum formation and marine mudstones of the f-ia unit (fjerritslev formation) in the southwestern and central parts of the basin (fig. 20, following page 498). towards the east, north-east and north, sandstone-dominated strata of the gassum formation gradually take over. on the skagerrak– kattegat platform and in terne-1 and hans-1 in the sorgenfrei–tornquist zone, the sequence consists exclusively of the gassum formation. the sequence spans the lower part of the o. aspinata zone of michelsen (1975, 1989a), and the lower parts of the d. priscum and pinuspollenites– trachysporites zones of dybkjær (1991) and poulsen (1992a, 1996). base of sequence, sb 9 and lowstand systems tract sb 9 is marked by a regional change from forestepping to backstepping sedimentary packages (fig. 20). in the stenlille area, sb 9 is marked by shoreface sandstones overlying heteroliths (fig. 16). on the thisted salt dome, fluvial sandstones cut down into marine mudstones and sandstones of the gassum formation. deep incision is also indicated by sb 9 in the himmerland graben, where cores and logs show a thick valley-fill of estuarine sandstones overlain by lagoonal siltstones and mudstones with abundant rootlets and a thin coal seam in års-1 (fig. 9b) and fluvial–estuarine sandstones in hyllebjerg-1 (fig. 20). farther to the north-east, sb 9 is traced into shoreface sandstones (vedsted-1) and to the base of fluvial sandstones, 2–10 m thick overlying marine mudstones and sandstones (flyvbjerg-1, frederikshavn-1, -2, -3). sandstones most likely of fluvial origin were cored in børglum-1 (fig. 5). the lowstand systems tract thus consists of fluvial, estuarine and shoreface sandstones passing basinwards into offshore heteroliths and mudstones. in most of the basin, sb 9 occurs within lowermost hettangian strata. towards the basin margins, it is correlated with the widespread base hettangian unconformity present on bornholm, in skåne, poland and germany (troedsson 1948; gry 1969; dadlez 1976; lund 1977; guy-ohlson 1981; gravesen et al. 1982; sivhed 1984; hallam 1988, 1992; bloos 1990; pieńkowski 1991; norling et al. 1993; ahlberg & arndorf 1994; surlyk et al. 1995). sb 9 thus records the greatest regional basinwards shift in facies of the upper rhaetian – lowermost hettangian, accompanied by fluvial incision and shoreface erosion. 493 transgressive systems tract the lowstand deposits are capped by ts 9 overlain by backstepping to aggrading, offshore heteroliths and mudstones, 3–25 m thick (figs 19, 20). cores from stenlille-5 show lowstand shoreface sandstones overlain by mudstones resting on a transgressive surface of marine erosion (fig. 16). in the southern part of the basin, ts 9 is marked by a change from marine heteroliths and silty mudstones to more uniform mudstones. in most of the basin, ts 9 coincides with the lithostratigraphic boundary between the gassum and fjerritslev formations. towards the north-east, ts 9 occurs within the gassum formation, and the transgressive systems tract consists of marine mudstones and 3–4 m of shoreface sandstones (e.g. flyvbjerg-1, vedsted-1; figs 19, 20). closer to the basin margin, the transgressive systems tract consists of aggrading coastal plain deposits with coal beds (e.g. hans-1; fig. 1). in the frederikshavn and skagen wells, c. 20 m of stacked lagoonal parasequences, indicated by serrated log-pattern and cores containing a mixed brackish palynomorph assemblage (dybkjær 1991, tables 3, 4), are overlain by offshore mudstones (fig. 15). in the gassum-1 well, the johnstoni subzone is indicated 14–15 m above ts 9 by the occurrence of curviceras sp. (fig. 13c; poulsen 1996), indicating together with the identified palynomorphs that ts 9 occurs in the lowermost hettangian. mfs 9 and highstand systems tract the maximum flooding surface mfs 9 is overlain by c. 25–75 m of marine mudstones primarily showing an aggrading pattern in the southern and central parts of the basin. the mudstones have a low content of ostracods, foraminifera and infaunal deposit-feeding bivalves and a high content of land-derived organic matter (michelsen 1975, pedersen 1986; dybkjær 1991). cores show poorly laminated to homogeneous, dark grey mudstones with scattered siltstone lenses and shell debris. farther to the north-east is a 60–75 m thick succession of aggrading shoreface and fluvial–estuarine sandstones interbedded with thin marine sandstones and mudstones (børglum-1, flyvbjerg-1; figs 5, 12). the mudstones sometimes contain a rich, but low-diversity ostracod fauna (fig. 12, core 7; michelsen 1975). biostratigraphic evidence suggests that mfs 9 belongs to the lower mid-hettangian, being located above the lowermost hettangian sb 9 and below the upper hettangian log marker a of michelsen (1989b). sequence fj 2 (upper hettangian – lowermost sinemurian) sequence fj 2 consists of marine mudstones and heteroliths of the upper f-ia unit (fjerritslev formation) in most of the basin (fig. 21, following page 498). along the north-eastern and eastern margins of the sorgenfrei– tornquist zone, the sequence consists of sandstones of the gassum formation. it spans the upper part of the o. aspinata zone (ostracods) of michelsen (1975, 1989a), part of the d. priscum zone (dinoflagellates), the upper part of the pinuspollenites–trachysporites zone and the lowermost part of the cerebropollenites macroverrucosus zone (spore–pollen) of dybkjær (1991) and poulsen (1992a; fig. 18). the sequence appears to be absent locally on the skagerrak–kattegat platform due to erosional truncation (fig. 20). base of sequence and lowstand systems tract the base of the sequence, sb 10, is defined by marine shoreface sandstones, heteroliths or sandy mudstones constituting the lowstand systems tract sharply overlying uniform to weakly forestepping marine mudstones in the fjerritslev trough, himmerland graben and at stenlille (fig. 21). farther into the basin, sb 10 appears to be conformable and is marked by a subtle change to more silty mudstones forming a thin lowstand systems tract. towards the north-eastern margin of the fjerritslev trough, sb 10 is tentatively traced into aggrading marine and fluvial sandstones along the børglum fault (børglum-1, flyvbjerg-1; fig. 20). sb 10 probably belongs to the upper hettangian angulata zone as it is located between log markers a and b of michelsen (1989b; fig. 21). the lowstand systems tract is capped by ts 10. transgressive systems tract, mfs 10, highstand and forced regressive systems tracts the transgressive systems tract overlying ts 10 consists of 10–20 m of aggrading to slightly backstepping marine mudstones in most of the basin passing into backstepping marine shoreface and offshore sandstones and heteroliths, 20–30 m thick close to the børglum fault. in the basinal mudstone-dominated successions, ts 10 corresponds to log marker b of michelsen (1989b; fig. 21). the maximum flooding surface is well-defined in the fjerritslev trough and himmerland graben, where it 494 is overlain by distinctly forestepping offshore mudstones that form a c. 8–15 m thick highstand systems tract. distally in the basin, mfs 10 is positioned within relatively homogeneous mudstones. the highstand deposits in the fjerritslev trough are sharply overlain by weakly forestepping to aggrading shoreface sandstones, c. 5–20 m thick that are interpreted as a forced regressive systems tract that wedges out towards the deeper parts of the basin (figs 8, 12, 20). mfs 10 is located below the f-ia/f-ib boundary in the lower sinemurian bucklandi zone and above log marker b which was suggested to occur at the hettangian–sinemurian boundary by michelsen (1989b). however, in stenlille-2, mfs 10 seems to be located in the hettangian pinuspollenites–trachysporites zone (spore–pollen; dybkjær 1991). in sæby-1, it occurs at the top of the hettangian defined by miospores (church et al. 1986; michelsen & nielsen 1991). a latest hettangian to earliest sinemurian age, most likely the angulata zone, is thus suggested for mfs 10. sequence fj 3 (lower–upper sinemurian) sequence fj 3 consists of the topmost part of the f-ia unit and the lower f-ib unit in most of the basin. in the sorgenfrei–tornquist zone and on the skagerrak–kattegat platform, the sequence includes the uppermost part of the gassum formation and the f-i member (frederikshavn wells; fig. 20; nielsen & japsen 1991). based on ostracods, the sequence spans the uppermost part of the o. aspinata zone, the cristacythere betzi – cristacythere crassireticulata zone, and the lower part of the ogmoconchella danica zone, including the progonoidea reticulata subzone of michelsen (1975, 1989b). it spans the upper part of the d. priscum and the lower part of the liasidium variabile zones (dinoflagellates), and the c. macroverrucosus zone (spore– pollen) of dybkjær (1988, 1991) and poulsen (1992a, 1996). base of sequence, lowstand systems tract and ts 11 the base of the sequence (sb 11) is defined by small but distinct log changes indicating a relatively abrupt change from uniform mudstones to heterolithic mudstones in most of the basin, and to coarser heteroliths in the himmerland graben (figs 20, 21). the change is interpreted to reflect shallowing and defines the conformable sb 11 overlain by lowstand heteroliths. sb 11 occurs in the upper part of the o. aspinata zone above log marker b of michelsen (1989b) and just below the top of the f-ia unit, suggesting that it occurs in the bucklandi zone. in the stenlille area, the lowstand systems tract is 10–15 m thick with shoreface sandstones sharply overlying marine mudstones (fig. 17). from the basin, sb 11 is traced into dominantly aggrading parasequences of marine shoreface sandstones and thin mudstones in the vedsted-1 well in the fjerritslev trough (fig. 20). the aggrading parasequences constituting the lowstand systems tract are capped by a transgressive surface ts 11 overlain by two backstepping shoreface parasequences. sb 11 and ts 11 are traced further to flyvbjerg-1 and børglum-1, where the surfaces tend to amalgamate on top of a 12–20 m thick unit of marine shoreface sandstones, interpreted as the highstand or forced regressive systems tract of the underlying sequence. the logs from the three closely spaced frederikshavn wells indicate 4–24 m of fluvial sandstones overlying sb 11 that shows variable depths of incision. the fluvial sandstones belong to the lowstand systems tract and are overlain by marine transgressive sandstones that yield d. priscum dinoflagellate cysts (figs 15, 20; dybkjær 1991). ts 11 is thus no younger than the bucklandi zone, as d. priscum cysts disappear or occur only very sporadically above this zone (woollam & riding 1983; poulsen & riding 2003, this volume). ts 11 is traced further to skagen-2, where cores with poor recovery show a few metres of siltstones with the marine bivalve oxytoma sinemuriensis indicating the bucklandi to semicostatum zones above ts 11 (sorgenfrei & buch 1964). core samples indicate that the transition from the hettangian pinuspollenites– trachysporites zone to the sinemurian c. macroverrucosus zone (poulsen 1992a) occurs very close to ts 11, which coincides with the boundary of the gassum and fjerritslev formations. farther to the south, in the sorgenfrei–tornquist zone, ts 11 also coincides with the boundary between the gassum and fjerritslev formations (terne-1 and hans-1) which occurs just above the hettangian–sinemurian boundary (michelsen & nielsen 1991). in the øresund-8 and -9 wells, sandy paralic deposits without ostracods are overlain by transgressive marine mudstones with ostracods of the c. betzi – c. crassireticulata zone corresponding to the semicostatum zone (michelsen 1975). in the basinal successions of continuous mudstones, ts 11 coincides with the top of the f-ia unit, where subtle log changes indicate a shift from lowstand heterolithic mudstones to transgressive homogeneous mudstones (fig. 21). ts 11 typically occurs 10–20 m 495 below log marker c, which michelsen (1989b) refers to the boundary between the bucklandi and semicostatum zones. hence, based on the ostracod and palynomorph data, the significant flooding represented by ts 11 occurred in the latter part of the bucklandi zone and appears to be of the same age in the entire basin. the flooding caused a significant decrease in diversity and density of the ostracod and bivalve faunas with a marked low in the semicostatum zone (pedersen 1986; michelsen 1989b). lithostratigraphically, the transgressive surface corresponds to the f-ia/f-ib boundary in most of the basin and to the gassum/fjerritslev formation boundary in the sorgenfrei–tornquist zone and on the skagerrak–kattegat platform. this indicates a slightly older age for the formation boundary than that suggested by michelsen (1975, 1978), who based the younger age on a few ostracod fragments from cuttings and poor core samples from børglum-1, flyvbjerg-1 and frederikshavn-2 that were possibly contaminated by drilling mud (o. michelsen, personal communication 1994). transgressive systems tract, mfs 11 and highstand systems tract in the central parts of the basin, the transgressive systems tract consists of a succession of very homogenous marine mudstones, up to 150 m thick, showing a subtle backstepping pattern overlain by weakly forestepping mudstones of the highstand systems tract (figs 20, 21). mfs 11 occurs between the log markers d and e, which are referred to the turneri and obtusum zones, respectively (michelsen 1989b). a very pronounced thinning of both systems tracts is indicated towards the north-east, where they thin to c. 25 m in børglum-1 and to 10–15 m in the frederikshavn wells; in skagen-2, they form an amalgamated succession of siltstones, less than 10 m thick. this dramatic thinning of the sequence is supported by biostratigraphic evidence indicating that the sinemurian is very thin in the børglum-1, frederikshavn-2, skagen-2 and sæby-1 wells (michelsen 1975; dybkjær 1991; michelsen & nielsen 1991; poulsen 1992a). much of the thinning seems to be primary, as deep erosion below sb 12 cannot be demonstrated. sequence fj 4 (upper sinemurian – lower pliensbachian) sequence fj 4 consists of marine mudstones of the f-ib unit in most of the basin and includes the lower part of the f-ii member on the skagerrak–kattegat platform (figs 20, 21). it spans part of the o. danica and g. apostolescui – k. (k.) foveolata zones (ostracods) of michelsen (1975, 1989b), and parts of the l. variabile zone (dinoflagellates) and the c. macroverrucosus zone (spore–pollen) of dybkjær (1991) and poulsen (1996). base of sequence, lowstand and transgressive systems tracts, and mfs 12 the base of the sequence, sb 12, is marked by subtle log changes, indicating that homogeneous mudstones are conformably overlain by slightly coarser mudstones over much of the basin including the sorgenfrei–tornquist zone (figs 20, 21). sb 12 is more clearly marked towards the north-east, where it is overlain by heterolithic sandstones (børglum-1, sæby-1 and frederikshavn wells). sb 12 coincides with log marker g, referred to the oxynotum zone by michelsen (1989b). in the basinal parts, the lowstand deposits consist of marine mudstones, less than 5 m thick, that coarsen to siltstones and bioturbated, heterolithic lower shoreface sandstones towards the north-east (børglum-1 and frederikshavn wells; fig. 5). the lowstand systems tract is overlain by uniform, aggrading to weakly backstepping marine mudstones capped by mfs 12 (fig. 20). mfs 12 coincides with log marker h (fig. 21) dated to the raricostatum zone and corresponds to a very marked low in the density and diversity of ostracods and bivalves in cores from gassum-1 (pedersen 1986; michelsen 1989b). a restricted shale facies without infaunal suspension feeders dominates at this level in the gassum cores (pedersen 1986). mfs 12 can be traced to most well sections including wells on the skagerrak–kattegat platform. highstand systems tract the highstand systems tract is 50–70 m thick in the sæby-1 and frederikshavn wells, and shows an aggrading to forestepping pattern (fig. 20). although recovery was poor, the cores from frederikshavn-1 and -2 exhibit silty and sandy, bioturbated mudstones with shell debris, that coarsen upwards to muddy, bioturbated, fine-grained sandstones also with shell debris, interpreted as the deposits of the offshore – lower shoreface transition zone. the well logs supported by cores indicate that the upper 15 m of the highstand systems tract consist of prograding lower shoreface sandstones 496 (fig. 22, following page 498). the lower part of the lower pliensbachian g. apostolescui – k. (k.) foveolata subzone of michelsen (1975) is indicated in core 5 just above mfs 12 in frederikshavn-2. the same zone occurs in skagen-2 in the upper 3 m of the highstand systems tract, which consists of 5–10 m of intensely burrowed muddy siltstones and fine-grained, lower shoreface sandstones with abundant bivalve fragments. thus, the highstand deposits belong to the lower lower pliensbachian. the palynomorph assemblage resembles that of the contemporaneous shoreface sandstones of the hasle formation on bornholm in that it includes mendicodinium reticulatum, tasmanites sp. and various acritarchs (poulsen 1992a; koppelhus & nielsen 1994). sequence fj 5 (lower–upper pliensbachian) sequence fj 5 consists of the main part of the f-ii member (f-iia and most of the f-iib beds of michelsen (1989a) where f-ii is subdivided) over much of the basin including the sorgenfrei–tornquist zone (fig. 21). north-east of the børglum fault, the sequence consists of the uppermost part of the f-ii member and the lowermost part of the f-iii member. with reference to the ostracod zonation, the sequence spans the upper part of the g. apostolescui – k. (k.) foveolata subzone and the lowermost part of the o. adenticulata – n. (n.) simplex zone (fig. 18) of michelsen (1975, 1989a). it spans the upper part of the l. variabile zone and the lower part of the l. spinosa zone (dinoflagellates), and the upper part of the c. macroverrucosus zone (spore– pollen) of dybkjær (1991) and poulsen (1996). base of sequence, lowstand and transgressive systems tracts the lower sequence boundary (sb 13) is marked by an abrupt log-break on sonic and resistivity logs, indicating homogenous marine mudstones overlain by a succession of silty and sandy mudstones that is typically 1–10 m thick although locally increasing to more than 40 m (fjerritslev-2). south-west of the børglum fault, sb 13 coincides with the base of the f-ii member, and tracing of log-patterns from hyllebjerg-1 to wells located on salt structures indicates that sb 13 is conformably developed (fig. 21). the silty and sandy mudstones are overlain by transgressive mudstones and are interpreted as a lowstand systems tract which corresponds to the f-iia beds of michelsen (1989b). a core just above the base of the f-ii member in gassum-1 shows a maximum in density and diversity of ostracods indicating a well-oxygenated environment (michelsen 1989b). the base of the f-ii member occurs in the lower part of the ibex zone, and is interpreted to be isochronous in most of the basin. north-east of the børglum fault, sb 13 occurs in the upper part of the f-ii member and the lowstand systems tract consists of 10–20 m of aggrading to weakly backstepping marine sandstones (sæby-1 and frederikshavn wells; fig. 22, see also fig. 25). these are sharply overlain by marine sandstones and mudstones, 6–8 m thick, constituting the transgressive systems tract (fig. 15). the mudstones contain a relatively rich and diverse ostracod fauna of the g. apostolescui – k. (k.) foveolata subzone, including both the index fossils and o. mouhersenis and o. amalthei, which indicate the upper lower pliensbachian (michelsen 1975; o. michelsen, personal communications 1991, 1993). a hiatus corresponding to the margaritatus zone and the lower part of the spinatum zone occurs in the frederikshavn cores, and the transgressive systems tract is truncated by sb 15 (figs 15, 22). mudstones of the transgressive systems tract in both the frederikshavn-2 and skagen-2 wells contain dinoflagellate cysts indicative of the nannoceratopsis gracilis zone (dybkjær 1991; poulsen 1992a); integration of the ostracod and dinoflagellate data indicates the middle to upper davoei zone (fig. 18). thus ts 13 is probably situated in the ibex zone or in the lower davoei zone, as it is located just above sb 13. mfs 13, highstand systems tract and thickness variations mfs 13 is relatively clearly marked in the himmerland graben, where the sequence is 30–60 m thick, and in the fjerritslev trough, where the thickness increases to c. 100 m in the fjerritslev-2 well, located close to the bounding fault. mfs 13 can be traced to the eastern part of the fjerritslev trough (børglum-1, flyvbjerg-1), but is not present east of the børglum fault (fig. 22, see also fig. 25). the sequence thins over salt structures with a minimum thickness of less than 5 m in kvols-1, and it appears that only the part of the sequence close to the point of maximum flooding is represented (fig. 21). the influence of the salt structures and the coarser grain size compared to sequences fj 1–4 suggest a shallower water depth during deposition compared to the underlying sequences. mfs 13 occurs in the upper lower pliensbachian in the børglum-1, f-1, fjerritslev-2, hyllebjerg-1 and 497 rønde-1 wells. in oddesund-1 and rødding-1, it occurs just below the top of the lower pliensbachian. in flyvbjerg-1, mfs 13 is identified below the base of the upper pliensbachian. since the available evidence indicates that the transgressive systems tract in frederikshavn-2 belongs to the davoei zone, mfs 13 is referred to the davoei zone, where both the diversity and density of ostracods are low (michelsen 1989b). sequence fj 6 (upper pliensbachian) sequence fj 6 consists of the upper f-ii member (upper f-iib and f-iic) and the lower f-iii member in the basin including the sorgenfrei–tornquist zone (figs 21, 22). the sequence spans the ostracod o. adenticulata – n. (n.) simplex zone of michelsen (1975, 1989a) with the exception of the lowermost and uppermost parts. it is referred to the l. spinosa zone of poulsen (1996). sequence fj 6 is not present on the skagerrak–kattegat platform, where a hiatus corresponding to the margaritatus zone and part of the spinatum zone is indicated, as discussed under sequence fj 7. base of sequence, lowstand and transgressive systems tracts, and mfs 14 the base of the sequence is marked by a subtle change from forestepping mudstones to backstepping silty and sandy mudstones across the conformable sb 14 in the middle part of f-iib in the himmerland graben. sb 14 can be traced to børglum-1 and flyvbjerg-1, where it is marked by a subtle change from forestepping to backstepping sandy mudstones and muddy sandstones (fig. 22). the log change is more pronounced in wells located on salt structures, suggesting that sb 14 is sharply developed (e.g. oddesund-1, rødding-1, skive-1; fig. 21). a hiatus occurs in the upper part of the f-ii member in these wells (michelsen 1989b) indicating that non-deposition or submarine erosion prevailed on the structures during the formation of the sequence boundary. sb 14 is located in the margaritatus zone, occurring above the top of the lower pliensbachian and below log marker i, which was referred to the margaritatus zone by michelsen (1989b). the lowstand systems tract is 15–40 m thick, and logmotifs and cuttings samples indicate that it consists of sandy and silty mudstones and fine-grained, heterolithic sandstones which show an aggrading to weakly backstepping pattern (figs 21, 22). it is sharply overlain by mudstones of the transgressive systems tract, which shows very smooth log-patterns in the lower part in most wells indicating 5–20 m of homogeneous marine mudstones. these are overlain by more variable mudstones showing both coarseningand fining-upwards trends. however, the overall backstepping culminates at mfs 14, just above log marker j of michelsen (1989b), which occurs close to the boundary between the margaritatus and spinatum zones. the transgressive systems tract ranges in thickness from c. 10 m in børglum-1 to more than 60 m in basinal sections. a core from flyvbjerg-1 illustrates parts of two coarsening-upwards parasequences of shell-bearing mudstones and intensely bioturbated muddy, fine-grained sandstones with erosion surfaces mantled by thin shell and pebble layers. highstand systems tract the highstand systems tract is up to c. 40 m thick and typically consists of weakly forestepping mudstones. in the eastern part of the fjerritslev trough, it consists of a succession of distinctly forestepping sandstones and mudstones, 20–30 m thick (fig. 22). sequence fj 7 (upper pliensbachian – middle toarcian) sequence fj 7 consists of the major part of the f-iii member over most of the basin; in the eastern part of the fjerritslev trough and on the skagerrak–kattegat platform, it includes the lowermost part of f-iv member. the lower part of the sequence corresponds to the uppermost part of the ostracod o. adenticulata – n. (n.) simplex zone. this zone is overlain by a thick section that is almost barren of ostracods, providing only weak evidence of the toarcian stage (michelsen 1975, 1989a, b). with reference to dinoflagellates, the sequence spans the l. spinosa subzone b and the m. semitabulatum subzone of poulsen (1996). base of sequence and lowstand systems tract the base of the sequence (sb 15) is clearly marked by a succession of marine shoreface sandstones, 5–12 m thick, overlying a regressive surface of marine erosion in the børglum-1, flyvbjerg-1, haldager-1 and vedsted-1 wells (figs 22, 23). in the himmerland graben, sb 15 is marked by a change from marine mudstones to a succession of silty or sandy mudstones, up to 25 m thick, 498 most clearly shown by the sonic and resistivity logs (fig. 22). on the skagerrak–kattegat platform, sb 15 is defined at the base of a very distinct peak on the gammaray log and the sp-log in the sæby-1, frederikshavn-1, -2, -3, and skagen-2 wells. this peak was cored in frederikshavn-2, where it is marked by less than 1 m of calcite-cemented marine sandstone bounded by mudstones (figs 15, 22). the mudstones below the sandstone bed contain a rich ostracod fauna typical of the lower pliensbachian g. apostolescui – k. (k.) foveolata subzone and belong to sequence fj 5, whereas the mudstones above the sandstone contain a fauna typical of the upper part of the o. adenticulata – n. (n.) simplex zone, belonging to the spinatum zone (michelsen 1975, 1989b; o. michelsen, personal communication 1993). this implies a significant upper pliensbachian hiatus at the base of the sandstone corresponding to sequence fj 6 and the upper part of fj 5. no evidence of subaerial exposure has been observed in the mudstones below sb 15, suggesting either that the missing section was removed by submarine erosion or that a subaerial erosion surface was obliterated by transgressive erosion. sb 15 coincides with log marker k of michelsen (1989b; figs 21, 22), and as it occurs 1–20 m below the top of the upper pliensbachian, it is well-dated to the upper part of the spinatum zone. the shoreface sandstones and their basinwards silty mudstone equivalent constitute the lowstand systems tract, which is capped abruptly by ts 15 in the fjerritslev trough. ts 15 is dated to the pliensbachian–toarcian boundary as it coincides with, or occurs very close to, the upper boundary of the ostracod o. adenticulata – n. (n.) simplex zone in most well sections. core samples from the lowstand systems tract in fjerritslev-2 and sæby-1 indicate the upper pliensbachian – lowermost toarcian subzone b of the dinoflagellate l. spinosa zone (forbes et al. 1985; poulsen 1992a, 1996). in all the investigated well sections, l. spinosa is absent in the succeeding transgressive systems tract. in anholt-4, l. spinosa cysts are only present in upper pliensbachian silty sands at the base of the borehole (seidenkrantz et al. 1993), and disappear a few metres below a transgressive surface overlain by marine clay. it is thus likely that the lowermost part of the anholt-4 section belongs to the lowstand systems tract, and the transgressive surface observed in this well correlates with ts 15. the anholt-4 section is thus considered to record the same overall development during the late pliensbachian – early toarcian as the other well sections in the area, contradicting the interpretation of seidenkrantz et al. (1993; see also nielsen et al. 2003, this volume). transgressive systems tract, mfs 15 and highstand systems tract the transgressive systems tract comprises a well-developed backstepping succession of marine mudstones and heteroliths, 15–40 m thick in most wells (figs 21, 23). the mudstones are almost barren of ostracods except 499 gamma-ray sp res sorgenfrei–tornquist zone central part haldager-1 ts 22 sb 22 sb 21 mfs 20 sb 20 sb 19 ts19 mfs 18 sb 18 mfs 17 sb 17 mfs 16 sb 16 mfs 15 fs ts 15 sb 15 m. semitabulatum zone l. spinosa subzone o. adenticulata – n. (n.) simplex zone p. na su ta z on e 10 0 m fl yv bj er g fm h al da ge r sa nd f or m at io n fj er ri ts le v fo rm at io n fl 1 ha 3 ha 2 ha 1 fj 6 fj 7 fj 8 fj 9 fj 10 fig. 23. well logs from the upper part of the fjerritslev formation, the haldager sand formation and the lowermost part of the flyvbjerg formation in the haldager-1 well showing sequences, key surfaces and the position of cores. biozones are based on michelsen (1975) and poulsen (1992a). note that the top of the dinoflagellate p. nasuta subzone is defined by a core sample of marine mudstones of the flooding surface that caps the lowermost parasequence overlying sb 19; this sequence boundary corresponds to the regional ‘base middle jurassic unconformity’, here developed as a regressive surface of marine erosion overlain by a lst of marine shoreface sandstones. for pyritic casts, and contain abundant spheripollenites psilatus together with amorphous organic matter implying reducing conditions (michelsen 1975; dybkjær 1991). the ostracods disappear just below ts 15 in the fjerritslev-2, hyllebjerg-1 and års-1 wells, which are located in the deep parts of the basin, whereas they disappear just above ts 15 in the frederikshavn-1 and -2 wells located on the shallow skagerrak–kattegat platform (michelsen 1975, 1989a, b). the disappearance of the benthic ostracods thus seems to be related to the development of anoxia associated with the transgression that formed ts 15. within the transgressive systems tract in vedsted-1 and haldager-1, a flooding surface is overlain by prograding units. the flooding surface corresponds to log marker l of michelsen (1989b), which possibly occurs in the tenuicostatum zone. the marker is traced to other well sections, but the prograding pattern is not recognised elsewhere. mfs 15 is distinctly marked by high gamma-ray, very low sonic and low resistivity values and rightwards deflection on the sp-log; severe caving of the condensed section results in anomalous values (i.e. low gamma-ray) in some wells (e.g. hyllebjerg-1; fig. 21). mfs 15 occurs just below log marker m in most wells, but is slightly above the marker in kvols-1. the log marker is correlated to the falciferum zone (michelsen 1989b). core pieces of greyish-green mudstones from c. 10 m below mfs 15 in haldager-1 and c. 10 m above mfs 15 in flyvbjerg-1 yield dinoflagellates of the m. semitabulatum subzone, which ranges from the uppermost tenuicostatum zone to the lower bifrons zone (poulsen 1992a, 1996); this supports a falciferum zone age for mfs 15. the highstand systems tract mainly consists of a succession of marine mudstones, up to c. 125 m thick in fjerritslev-2, showing a weak coarsening-upwards to siltstones in the uppermost part. on the skagerrak–kattegat platform, the highstand deposits contain progradational units of coarser grained mudstones. sequences fj 8, fj 9 and fj 10 (middle toarcian – lower aalenian) the succession referred to sequences fj 8, fj 9 and fj 10 consists of marine mudstones and sandstones of the f-iv member. its top corresponds to the ‘base middle jurassic unconformity’. the succession is almost barren of ostracods. cores from the vedsted-1 and haldager-1 wells have yielded a monospecific fauna with poor stratigraphic significance, but suggesting a broad toarcian– aalenian age (michelsen 1975, 1978, 1989b). a slightly more diverse fauna from the øresund-2 well, supported by foraminifers, suggests an aalenian age (michelsen 1975). dinoflagellate cysts of the parvocysta nasuta subzone have been reported from the farsø-1, fjerritslev-2, haldager-1, lavø-1, terne-1, vedsted-1 and øresund-2 wells, indicating the middle toarcian – lower aalenian (forbes et al. 1985; poulsen 1992a, 1996). the succession is thickest in the deep part of the sorgenfrei–tornquist zone, and also occurs in the himmerland graben and on the skagerrak–kattegat platform; only the lower part of the succession is present in the k-1, kvols-1, mors-1 and rønde-1 wells, and it is absent in the southern and south-easternmost part of the basin. in haldager-1, three sharp-based sandstones up to 10 m thick occur within forestepping and backstepping marine mudstones (figs 23, 24; fig. 24 follows page 498). the sandstones are interpreted as shoreface sandstones overlying regressive surfaces of marine erosion implying that lowstand shorelines were established in the sorgenfrei–tornquist zone during the middle–late toarcian. the bases of the sandstones are interpreted as sequence boundaries, sb 16–18, which are traced to terne-1, located on strike towards the south-east. updip from haldager-1, the sequence boundaries appear to be marked by thin veneers of sandstone, possibly reflecting lowstand erosion and bypass and subsequent transgressive erosion (e.g. farsø-1, frederikshavn-2, vedsted-1; fig. 24). thin lagoonal units with coals were cored in the skagen-2 well, in the uppermost part of the f-iv member. the member typically consists of marine mudstones and these unusual lagoonal deposits overlain by normal marine mudstones are interpreted as erosional remnants of transgressive, paralic deposits overlying sequence boundaries. the maximum flooding surfaces of the three sequences, mfs 16–18, are only well-defined in haldager-1 and terne-1. palynomorphs from core samples in haldager-1 show that the surfaces occur in the dinoflagellate p. nasuta subzone (poulsen 1992a, 1996). sb 16, sb 17 and mfs 16 can be traced to the himmerland graben (farsø-1) and updip to the skagerrak– kattegat platform, but the biostratigraphic resolution is too poor to confirm the correlations (fig. 24). sequences ha 1, ha 2, ha 3 and fl 1 (aalenian–oxfordian) the succession referred to sequences ha 1, ha 2, ha 3 and fl 1 consists of the haldager sand formation and the lower two-thirds of the flyvbjerg formation. it is 500 bounded by the ‘base middle jurassic unconformity’ below and sb 23 above (fig. 24). the succession is poorly dated as it primarily contains long-ranging spores and pollen, suggesting a broad middle – early late jurassic age. in the sorgenfrei–tornquist zone, however, palynomorphs indicating the aalenian–callovian and oxfordian are found (forbes et al. 1985; dybkjær 1991; michelsen & nielsen 1991; poulsen 1992a, b, 1996; seidenkrantz et al. 1993). the succession reaches a maximum thickness of c. 200 m in the sorgenfrei–tornquist zone decreasing to c. 50 m on the skagerrak–kattegat platform. in the north-west and central parts of the basin, the thickness is between 30–60 m depending on the position relative to salt structures and faults. the thickness decreases to c. 2–20 m in the southern and south-western part of the basin, and the succession is absent close to, and upon, the ringkøbing–fyn high. sequence ha 1 sequence ha 1 is only present in the deepest part of the sorgenfrei–tornquist zone. it wedges out towards the faulted margins of the fjerritslev trough and is absent in the børglum-1, fjerritslev-2, and flyvbjerg-1 wells (fig. 24). the base of the sequence (sb 19) is marked by a change from marine offshore mudstones of the fjerritslev formation to three shoreface sandstone units, c. 18–20 m thick, separated by fine-grained beds (haldager-1, terne-1; figs 23, 24). core fragments from the fine-grained beds in haldager-1 show a weakly bioturbated heterolith containing abundant plant material, pyrite and a marine palynomorph assemblage (ravn-sørensen 1989; poulsen 1992a). in terne-1, dinoflagellate cysts, tasmanites sp. and foraminiferal test linings occur (forbes et al. 1985). the biostratigraphic evidence indicates that sb 19 occurs in the lower aalenian opalinum zone, just below the base of the murchisonae zone, as suggested below. the sandstones are sharply overlain by transgressive mudstones and are referred to the lowstand systems tract. the upper sandstone in terne-1 is carbonate-cemented, a feature that is commonly seen at transgressive surfaces capping lowstand systems in the danish basin. the relatively gradual base of the sandstones in terne-1 and haldager-1 suggests that only limited shoreface erosion occurred in the deepest part of the sorgenfrei–tornquist zone during the sea-level fall that produced sb 19. in contrast, sb 19 is very sharply overlain by shoreface sandstones showing a weak overall forestepping progradational pattern in vedsted-1 (fig. 24). the overlying transgressive systems tract consists of backstepping parasequences with interbedded marine mudstones and sandstones (35 m thick in terne-1) overlain by forestepping parasequences defining the position of mfs 19. a similar pattern is absent in the haldager-1 and vedsted-1 wells, where only 6–9 m of mudstones and sandstones are preserved below sb 20, possibly reflecting a more shallow and proximal position than at terne-1. sequence ha 2 sequence ha 2 has only been identified in the deep part of the sorgenfrei–tornquist zone. the base of the sequence (sb 20) is overlain by a succession of sandstones with mudstone beds and thin coaly beds; this succession is c. 45 m thick in terne-1 and haldager-1 (figs 23, 24). the succession comprises a basal sandstone and three fining-upwards units overlain by a coarsening-upwards unit in haldager-1. a core from the upper fining-upwards unit consists of cross-bedded, fineto coarse-grained, slightly pebbly sandstones and laminated siltstones with large plant fragments. the succession is capped by laminated, silty mudstones with plant fragments and non-marine palynomorphs, as shown by core pieces (ravn-sørensen 1989). the succession is interpreted as a basal fluvial sandstone overlain by stacked, sandy estuarine channel and lagoonal units. the lower part of the succession in terne-1 shows a serrated log pattern which, supported by cuttings samples, indicates a thin fluvial sandstone overlain by interbedded lagoonal sandstones and mudstones with thin coaly seams and a mixed palynomorph assemblage including n. gracilis, tasmanites sp. and botryococcus sp. the upper part, showing a blocky log motif, consists of 18–19 m of barrier sandstones with thin coal seams and a mixed palynomorph assemblage. the successions, especially in terne-1, resemble a succession formed in a drowned estuary (allen & posamentier 1994; dalrymple et al. 1994), and are thus interpreted as estuarine valley-fills formed during rising sea level in the deep part of the sorgenfrei–tornquist zone. the terne-1 succession seems to reflect a slightly more seawards position than haldager-1. the valley-fills are overlain by weakly backstepping marine sandstones and mudstones, 26 m thick in terne-1, and aggrading to weakly prograding shoreface sandstones, 15–30 m thick and capped by thin transgressive mudstones in haldager-1 and vedsted-1 (fig. 24). the progradational pattern may reflect proximity to 501 sand-rich sources, while the terne-1 succession reflects a more seawards position. forestepping mudstones and sandstones of the highstand systems tract, 12 m thick, clearly define the position of mfs 20 in terne-1, while highstand deposits are thin or absent in vedsted-1 and haldager-1. the dinoflagellate cyst n. gracilis has its topmost occurrence in the highstand systems tract in terne-1 (forbes et al. 1985), suggesting that these deposits are not younger than early bajocian (poulsen & riding 2003, this volume). therefore, sb 20 was most likely formed in late aalenian – early bajocian times. sequence ha 3 sequence ha 3 is identified in the sorgenfrei–tornquist zone, the most basinwards part of the skagerrak–kattegat platform, in parts of the himmerland graben and in felicia-1 located just outside the fjerritslev trough (fig. 1). the sequence wedges out south-west of farsø-1 and is absent in hyllebjerg-1, where deep erosion is indicated below sb 22 (fig. 25, facing page 499). it thus oversteps the limits of sequences ha 1 and ha 2. in the deepest part of the sorgenfrei–tornquist zone, the base of the sequence (sb 21) is overlain by fining-upwards successions, 21–25 m thick, of coarseto fine-grained, muddy sandstones and thin mudstones, possibly with coal seams in the upper part (figs 23, 24). cuttings samples from terne-1 and cores from haldager-1 contain a mixed assemblage of marine and freshwater palynomorphs (forbes et al. 1985; ravn-sørensen 1989), and the cores comprise fineto medium-grained, laminated, weakly bioturbated sandstone with abundant organic debris, mudstone drapes and flakes. the fining-upwards successions are interpreted as estuarine channel-fills deposited during rising sea level. elsewhere, sequence ha 3 commences with fluvial sandstones, 7–25 m thick. the sandstones appear to occur in small channel units, 3–7 m thick, as suggested by subtle log fluctuations (fig. 24). the sandstones were cored in farsø-1 and frederikshavn-1 and are sharply capped by lacustrine mudstones (farsø-1) or lagoonal mudstones (frederikshavn wells) reflecting a further rise in sea level (fig. 10b). in flyvbjerg-1, sb 21 cuts more than 50 m into marine mudstones of the fjerritslev formation relative to the neighbouring børglum-1 well, and the overlying sandstone-dominated section is interpreted as an estuarine valley-fill (fig. 25). sequence ha-3 is absent in børglum-1 as indicated by log correlation and biostratigraphic evidence that show that the lower part of the middle jurassic is missing here (ravn-sørensen 1989). in felicia-1, the sequence consists of c. 17 m of carbonaceous sandstones and mudstones with shell fragments and a terrestrially dominated palynomorph assemblage, probably representing a transgressive lagoonal unit. highstand deposits seem to be absent in all well sections except for haldager-1, where 8 m of coarseningupwards sandstones overlying a thin mudstone are interpreted as a prograding shoreface. palynomorphs from felicia-1 indicate an early bathonian age (stratlab a.s. 1988), and in combination with weak biostratigraphic evidence from other wells, it is suggested that sb 21 was formed in late bajocian – early bathonian times. sequence fl 1 sequence fl 1 oversteps the limit of sequence ha 3 and is present over most of the basin from skagen-2 to the south-western part of the basin. it includes the topmost part of the haldager sand formation and the lower two-thirds of the flyvbjerg formation. it shows a pronounced north-eastwards thickening (fig. 25). base of sequence, sb 22, lowstand and transgressive systems tracts the base of the sequence (sb 22) is overlain by shoreface sandstones in the deep part of the sorgenfrei–tornquist zone. in the southern and south-western part of the basin, where marine mudstones of the fjerritslev formation are deeply truncated, sb 22 is overlain by thin and discontinuous sandstones interpreted as fluvial sandstones preserved in channel scours below the transgressive surface, ts 22 (e.g. horsens-1, mejrup-1, oddesund-1; fig. 25). fluvial sandstones overlie lacustrine–lagoonal mudstones of sequence ha 3 in the farsø-1, felicia-1, frederikshavn-2 and vedsted-1 wells. in skagen-2, fluvial sandstones overlie lagoonal mudstones included in the fjerritslev formation (fig. 24). the sandstones are 1–18 m thick and were cored in the farsø-1, frederikshavn-1, skagen-2, vedsted-1 and års-1 wells (fig. 10b). sb 22 cuts deeply into the fjerritslev formation in hyllebjerg-1 in contrast to the nearby farsø-1 and års-1 sections, where sequence ha 3 is preserved below sb 22. this suggests that hyllebjerg-1 lay structurally higher, as is also indicated by a primary thinning of the uppermost lower jurassic marine mudstones (f-iii, f-iv members). the shoreface and fluvial 502 sandstones belong to the lowstand systems tract and are erosionally or sharply overlain by transgressive lagoonal deposits or marine mudstones. the transgressive surface (ts 22) coincides with the boundary between the haldager sand and flyvbjerg formations in most of the basin. this boundary has been dated to occur just above the callovian–oxfordian boundary (poulsen 1996). biostratigraphic evidence of the callovian stage is poor in the danish basin, and a significant hiatus between the haldager sand and flyvbjerg formations was assumed by michelsen (1989a). the interpretation presented here indicates that the formation boundary corresponds only to a diastem caused by transgressive erosion. the transgressive systems tract shows great lateral variation in composition and thickness. in the himmerland graben, it is c. 5–15 m thick and cores (års-1) show lagoonal deposits consisting of burrowed and rooted, wavy to lenticular bedded, occasionally wave-rippled to small-scale hummocky cross-stratified heteroliths of very fine-grained sandstones, coarse-grained siltstones and claystones. cores from skagen-2 and vedsted-1 comprise bioturbated, parallel-laminated, cross-laminated and wave-rippled sandstones and heteroliths, 2–3 m thick, with roots, plant fragments and brackish-marine dinoflagellates. these deposits are interpreted as transgressive lagoonal beach deposits overlying a lagoonal ravinement surface (fig. 24). locally, the fluvial lowstand deposits are non-erosionally overlain by muddy siltstones with rootlets and thin coal seams, indicating a gradual and gentle flooding (frederikshavn-3). towards the southwestern part of the basin, the systems tract thins to few metres of mudstones of presumed lagoonal origin. in the sorgenfrei–tornquist zone and on the skagerrak–kattegat platform, the transgressive systems tract also includes backstepping marine deposits up to c. 30 m thick (fig. 24). cores from the børglum-1, flyvbjerg-1, frederikshavn-1, -3 and skagen-2 wells exhibit burrowed, lenticular-bedded heteroliths, laminated carbonaceous mudstones, cross-laminated siltstones and fine-grained sandstones; calcite-cemented sandstone beds with bivalve shells occur in places. the palynomorph assemblages are fairly rich in dinoflagellates showing an upwards increase in density and diversity (forbes et al. 1985; church et al. 1986; ravn-sørensen 1989; poulsen 1992a, b). dinoflagellate cysts of the dichadogonyaulax sellwoodi ‘group’ have their topmost occurrences in the transgressive systems tract in terne-1, indicating an age no younger than the early oxfordian densiplicatum zone (poulsen 1996). highstand systems tract the highstand deposits are very thin in the southwestern part of the basin, consisting of a few metres of mudstones (fig. 24). the thickness increases to 5–22 m in the sorgenfrei–tornquist zone and on the skagerrak– kattegat platform, where cores consists of carbonaceous mudstones and burrowed, heterolithic, fine-grained sandstones with bivalve shells and carbonate-cemented beds (børglum-1, frederikshavn-1, haldager-1, skagen-2). marine palynomorphs have been reported from haldager-1 and skagen-2 (ravn-sørensen 1989; poulsen 1992a, b). the highstand deposits are thus interpreted to represent a marine offshore to lower shoreface environment. the palynomorphs of the transgressive and highstand systems tracts in the hyllebjerg-1, skagen-2, terne-1, vedsted-1 and års-1 wells indicate a general oxfordian age, possibly with hiati or thin biozones (forbes et al. 1985; poulsen 1992a, 1996). the dinoflagellates in skagen-2 show that mfs 22 occurs in the upper oxfordian – lowermost kimmeridgian scriniodinium crystallinum zone (poulsen 1996; poulsen & riding 2003, this volume). the upper sequence boundary, sb 23 the upper boundary of sequence fl 1, sb 23, is defined at the base of a unit of shell-bearing marine shoreface siltstones and sandstones, 2–10 m thick, in the uppermost part of the flyvbjerg formation in the fjerritslev trough. on the skagerrak–kattegat platform, sb 23 is overlain by fluvial sandstones. these lowstand deposits are capped by a transgressive surface (ts 23) that is overlain by marine mudstones of the børglum formation in most of the basin, including the sorgenfrei–tornquist zone and the skagerrak–kattegat platform (fig. 24). in the himmerland graben and in felicia-1, sb 23 is marked by a change from forestepping to backstepping marine parasequences, and the boundary between the flyvbjerg and børglum formation appears to be conformable. the base of the børglum formation occurs just above the oxfordian–kimmeridgian boundary in most of the basin, with a weak younging trend towards the northeast (poulsen 1992a, b, 1996). sb 23 is thus dated to the uppermost oxfordian. discussion the sequence stratigraphic interpretation described in detail above and summarised in figures 25 and 26 pro503 vides the basis for an improved understanding of the late triassic – jurassic evolution of the danish basin, both in terms of intrinsic (i.e. intrabasinal) and extrinsic (regional or global) controlling factors. timing and cause of formation of the regional intra–aalenian unconformity the ‘base middle jurassic unconformity’ is the most significant unconformity within the upper triassic – jurassic succession in the danish basin and the fennoscandian border zone (figs 10b, 11b, 21, 24–26). in most of the basin, it is a pronounced seismic unconformity that shows progressively deeper truncation of the lower jurassic fjerritslev formation and older formations towards the ringkøbing–fyn high. the unconformity is also developed as a significant erosion surface on the skagerrak–kattegat platform, whereas it passes into a seismic conformity in the deep part of the sorgenfrei– tornquist zone. the overlying succession shows an upper aalenian – volgian onlap onto the unconformity surface with a pronounced younging of the onlap towards the ringkøbing–fyn high. jurassic deposits are eroded on the high, and lower cretaceous strata onlap eroded triassic deposits or precambrian crystalline basement. the ‘base middle jurassic unconformity’ thus subdivides the upper triassic – jurassic succession into two principal tectono-stratigraphic units. formation of the deep, erosional unconformity required major uplift of the ringkøbing–fyn high and large parts of the basin (fig. 27). the onlap of the overlying strata shows that accommodation space was subsequently created due to renewed subsidence, as the eustatic sea-level rise alone cannot account for the onlap pattern. the unconformity is most precisely dated in the sorgenfrei–tornquist zone, where it passes into a relatively conformable, regressive surface of marine erosion (sb 19). generally, biostratigraphic information from the toarcian–aalenian deposits is poor, reflecting the restricted nature of the basin at this time, but data from the anholt-4, børglum-1, fjerritslev-2, haldager-1, terne-1, vedsted-1 and øresund-2 wells indicate that deposition of marine mudstones and siltstones continued uninterrupted from early jurassic into aalenian times in the sorgenfrei–tornquist zone (michelsen 1975; forbes et al. 1985; ravn-sørensen 1989; dybkjær 1991; michelsen & nielsen 1991; poulsen 1992a, b, 1996; seidenkrantz et al. 1993). the top of the upper toarcian – lower aalenian p. nasuta zone (dinoflagellates) is found above sb 19 in core samples of mudstones in haldager-1 (poulsen 1992a, b, 1996) from the flooding surface capping the lowermost parasequence of the lowstand systems tract of sequence ha 1 (fig. 23). the top of the p. nasuta zone coincides with the opalinum– murchisonae zone boundary (poulsen & riding 2003, this volume), and as aalenian marine mudstones are identified below sb 19, it is most likely that sb 19 occurs in the lower aalenian opalinum zone, just below the base of the murchisonae zone. thus the age of the pronounced basinwards shift in facies across the ‘base middle jurassic unconformity’ in the sorgenfrei–tornquist zone corresponds to the major basinwards shift in facies at the ‘mid-cimmerian unconformity’ in the north sea (underhill & partington 1993, 1994) and the major 177 ma sequence boundary of haq et al. (1988). the regional ‘mid-cimmerian unconformity’ in the north sea has been interpreted to reflect uplift and erosion related to emplacement of a mantle plume (whiteman et al. 1975; hallam & sellwood 1976; eynon 1981; ziegler 1990; underhill & partington 1993, 1994), a major eustatic sea-level fall (haq et al. 1988), or local response to footwall uplift (barr 1987; badley et al. 1988). underhill & partington (1993, 1994) portrayed a roughly circular pattern of the subcrop stratigraphy of the unconformity c. 800 km in diameter and centred over the junction of the central graben, the moray firth and the viking graben. based on this pattern, they interpreted domal uplift and erosion followed by deflation and gradual onlap. however, in their contouring of the crucial subcrop stratigraphy, underhill & partington (1993, 1994) assumed that missing sections on highs such as the east shetland platform, the mid 504 facing page: fig. 26. schematic time-stratigraphic sw–ne section from the ringkøbing–fyn high (rkf) across the danish basin and the sorgenfrei–tornquist zone (stz) to the skagerrak–kattegat platform (skp) showing depositional sequences and environments, related to the chronostratigraphy and lithostratigraphy. sequence stratigraphic key surfaces are drawn at their most likely age: the diagram illustrates the deep erosional truncation of the pre-lower aalenian strata and upper jurassic onlap, younging towards the rkf. the ‘base cretaceous unconformity’ merges with the ‘base middle jurassic unconformity’ close to the high, and cretaceous strata overlie precambrian basement on the rkf. the numerous hiati on the skagerrak–kattegat platform illustrate the limited accommodation space here due to relatively slow subsidence in norian–callovian times. the duration of the middle jurassic hiati is poorly constrained. time-scale from gradstein et al. (1994). 505 sw ne ma dolomitic limestones fluvial sandstones alluvial conglomerates, sandstones sabkhas and lacustrine calcareous, evaporitic mudstones estuarine/lagoonal sandstones, heteroliths, mudstones, coal beds shoreface sandstones/siltstones offshore mudstones; occasionally sandy or silty lacustrine mudstones hiatus norian rhaetian hettangian sinemurian pliensbachian toarcian aalenian bajocian bathonian callovian oxfordian kimmeridgian volgian ryazanian chronostratigraphy sequence lithostratigraphy tectonics danish basin rkf stz skp fr 3 150.7 142.0 154.1 159.4 164.4 169.2 176.5 180.1 189.6 195.3 201.9 205.7 209.6 fr 2 fr 1 bø 1 fl 1 ha 3 ha 2 ha 1 fj 10 fj 9 fj 8 fj 7 fj 6 fj 5 fj 4 fj 3 fj 2 fj 1 ga 1 vi 1 frederikshavn fm børglum fm flyvbjerg fm haldager sand fm f-iv mb f-iii mb f-ii mb f-ib f-i mb f-ia gassum fm vinding fm sk od fj er ri ts le v fm renewed regional subsidence including ringkøbing– fyn high basin expansion fault-controlled subsidence of sorgenfrei– tornquist zone uplift of ringkøbing– fyn high, ne-tilting of basin thermal subsidence local faulting and halokinesis fjerritslev trough himmerland graben local incision erosion over saltstructures condensed section ‘base cret. unconf.’ ‘base middle jurassic unconformity’ ‘mid-cimmerian unconformity’ north sea high and the ringkøbing–fyn high reflect local erosion due to younger footwall uplift or other factors that overprint the regional pattern. based on this assumption, contours were connected across the ringkøbing–fyn high. the contours of the subcrop stratigraphy of the unconformity in the danish basin roughly follow the trend of the ringkøbing–fyn high (fig. 28a). the same pattern is probably present in the north german basin south of the high, although the ‘base cretaceous unconformity’ partly masks the pattern, as post-aalenian jurassic deposits are missing in the danish wells from this area. the onlap pattern also parallels the general trend of the ringkøbing–fyn high (fig. 28b) and, together with the subcrop contours, testifies that the high including its northern and southern flanks were uplifted, eroded and later submerged resulting in the regional intra-aalenian unconformity. the unconformity can probably be traced to the rønne graben, where lower aalenian(?) shoreface sandstones are unconformably overlain by aalenian(?)–bajocian fluvial conglomerates (koppelhus & nielsen 1994). the unconformity also occurs in skåne, where marine sedimentation continued until the end of the early jurassic followed by continental and paralic deposition in middle jurassic time reflecting tectonic uplift and erosion over much of southern sweden, with the exception of the south-western part of skåne, hanö bay and a narrow fault-controlled belt across central skåne (norling & bergström 1987; norling et al. 1993; erlström et al. 1997; ahlberg et al. 2003, this volume). the age of the unconformity is not precisely determined, but appears to occur in the aalenian. the present-day distribution of the jurassic deposits is strongly influenced by later inversion and 506 d ep th ( m ) 0 500 1000 1500 late triassic early jurassic middle jurassic late jurassic e.cret. time (ma) 220 210 200 190 180 170 160 150 140 ringkøbing–fyn high basin centre skagerrak–kattegat platform himmerland graben fjerritslev trough fig. 27. subsidence curves for five selected positions representing the ringkøbing–fyn high (ullerslev-1), the distal danish basin (mejrup-1), the himmerland graben (hyllebjerg-1), the fjerritslev trough (vedsted-1, haldager-1) and the skagerrak–kattegat platform (sæby-1). the curves are constructed by the use of the yükler 1d forward modelling program (yükler et al. 1978); this adopts specified, decompacted lithologies and the time of formation of each model event corresponds to the sequences described herein (timescale of gradstein et al. 1994). the latest event essentially corresponds to the kimmeridgian–ryazanian frederikshavn formation. water depth that varied between 0–50 m during deposition of the sequences was averaged to 0 m to facilitate modelling and to reduce the effects of rapid sea-level fluctuations. the lithology and thickness of eroded sequences was determined from palaeogeography and by comparison to preserved sections. the five curves emphasise the aalenian–callovian uplift of most of the area, with the exception of the fault-bounded sorgenfrei–tornquist zone (fjerritslev trough), where subsidence continued with a low rate in the middle jurassic. the curves also indicate that regional subsidence gradually resumed during the late jurassic – early cretaceous in most of the basin. 507 l. sin em . sinemurian toarcian u. si nem . norway bornholm bornholm skåne ? ? ? ? u. toarcian u. pliensbach. l. toarcian l. pliensbach. sinemurian hettangian u. triassic hettangian hettan gian 6°e 10°e 14°e 6°e 10°e 14°e 58°n 56°n 58°n 56°n a b ? ? ? ? kimm.–volg. bajoc.–bathon. bajocian oxf. oxfordian skåne norway base middle jurassic subcrop subcrop to the ‘base cretaceous unconformity’ fault well intra-aalenian basinward shift in facies middle–upper jurassic absent subcrop to the ‘base middle jurassic unconformity’ 100 km the netherlands the netherlands germany germany onlap contours uncertain fault well intra-aalenian basinward shift in facies middle–upper jurassic absent onlap contours 100 km onlap on to the base middle jurassic unconformity fig. 28. a: generalised subcrop contour map of the ‘base middle jurassic unconformity’. the contours are based on the recognised sequences, lithostratigraphy from nielsen & japsen (1991), age determinations from dybkjær (1991), poulsen (1996), norling & bergström (1987) and erlström et al. (1997). sinemurian (sinem.) deposits are preserved ne of the sorgenfrei–tornquist zone in skåne, while toarcian deposits occur in hanö bay. b: generalised contour map showing the onlap to the ‘base middle jurassic unconformity’. bajocian deposits occur in hanö bay. both the subcrop contours (a) and the onlap contours (b) parallel the general trend of the ringkøbing–fyn high. bajoc., bajocian; bathon., bathonian; oxf., oxfordian; kimm., kimmerigian; volg., volgian. 508 509 erosion, but sedimentation was probably most continuous in the sorgenfrei–tornquist zone (norling & bergström 1987; erlström et al. 1997). basalts were intruded along nw–se-trending faults and fracture zones; the oldest basalts are palaeomagnetically dated to the toarcian–aalenian and radiometrically dated to the bajocian (167 ma), while related tuffites are dated to the aalenian by palynology (printzlau & larsen 1972; tralau 1973; klingspor 1976; norling & bergström 1987; bylund & halvorsen 1993). the roughly circular pattern of the subcrop contours shown by underhill & partington (1993, 1994) thus conthis and facing page: fig. 29. comparison of third-order sequence stratigraphic key surfaces from the danish basin and the fennoscandian border zone with ‘global’ surfaces recognised by haq et al. (1988), north sea surfaces documented by partington et al. (1993) and underhill & partington (1994), those from poland by feldman-olszewska (1997a, b), and from the uk by hesselbo & jenkyns (1998). timescale from gradstein et al. (1994); note that the numbers attached to the haq et al. (1988) surfaces relate to the original ages ascribed by these workers, included here solely for ease of reference. flicts with data from the danish basin and fennoscandian border zone. their model underestimates the areal effect of the north sea domal uplift and overlooks the possibility of at least one more dome centred in skåne. rather than suggesting a major circular to elliptical north sea dome, the contours suggest a broad uplifted area, as has also been suggested by ziegler (1982, 1990). in addition, the role of the long-lived sorgenfrei– tornquist zone has to be considered, as this fundamental fracture zone forms a buffer zone between the danish basin and the baltic shield that has accommodated late palaeozoic – mesozoic transtension, has experienced phases of volcanic activity (late carboniferous – early permian, middle jurassic, cretaceous) and has undergone late cretaceous – early cenozoic transpression resulting in tectonic inversion. contemporaneous with the regional aalenian–bajocian uplift, slow subsidence occurred in the sorgenfrei–tornquist zone and this zone of crustal weakness probably functioned as a hinge zone between the north-eastwards tilting basin and the baltic shield. thus, a model explaining both north sea uplift and volcanism, uplift of the danish basin and the ringkøbing–fyn high, and uplift and volcanism in skåne probably has to invoke lithospheric processes on a larger scale than a single transient plume. comparison of the sequence stratigraphic key surfaces to other schemes the third-order sequence stratigraphic key surfaces from the danish basin and the fennoscandian border zone can be compared and contrasted with those presented by haq et al. (1988), partington et al. (1993; underhill & partington 1994), feldman-olszewska (1997a, b) and hesselbo & jenkyns (1998; fig. 29). the haq et al. (1988) chart is claimed to portray a global, eustatic signal, while the compilations of partington et al. (1993) and underhill & partington (1994) are based on data from the large north sea basin, of which the danish basin formed a part. hesselbo & jenkyns (1998) compared the lower jurassic sequence stratigraphies of several british basins whereas feldmanolszewska (1997a, b) described surfaces of maximum regression and maximum transgression from the polish basin. the potential error of the assigned ages of the upper triassic – lower jurassic and upper jurassic surfaces in the danish area may amount to more than one standard zone in some cases. the number of upper triassic – lower jurassic sequences in the danish area equals that of haq et al. (1988), whereas partington et al. (1993) and underhill & partington (1994) recognised many fewer sequences. although the most recent, well-constrained time-scale (gradstein et al. 1994) is adopted here, the haq et al. (1988) surfaces are referred to according to their original age assignments. thus the mfs originally dated at 223 ma is abbreviated here to ‘mfs 223’, although its position has been re-calibrated to gradstein et al. (1994) on figure 29. partington et al. (1993) utilised maximum flooding surfaces to bound their sequences; these are numbered sequentially (j16 etc.). a detailed comparison indicates differences in the poorly dated upper triassic, where the ‘mfs 223’ of haq et al. (1988) occurs at the carnian/norian boundary, while mfs 1 probably occurs in the upper norian, and sb 5 in the early rhaetian where partington et al. (1993) indicate a basinwards shift in facies. in contrast, mfs 7 is well-dated to the uppermost rhaetian by ostracods and palynomorphs supported by ammonites, and is interpreted to correspond to ‘mfs 211.5’ of haq et al. (1988) and j02 of partington et al. (1993). sb 9 occurs in the lowermost hettangian, and is correlated with the regional significant base hettangian unconformity at the basin margins in skåne, bornholm, poland and germany (troedsson 1948; gry 1969; dadlez 1976; lund 1977; guy-ohlson 1981; gravesen et al. 1982; sivhed 1984; hallam 1988, 1992, 1997; bloos 1990; pieńkowski 1991; norling et al. 1993; ahlberg & arndorf 1994; surlyk et al. 1995; feldman-olszewska 1997a), whereas haq et al. (1988) positioned their ‘sb 211’ in the uppermost rhaetian. mfs 9 and ‘mfs 205.5’ of haq et al. (1988) indicate an early–middle hettangian maximum flooding in accordance with hesselbo & jenkyns (1998), whereas partington et al. (1993) indicate a regression. mfs 10 is dated to the uppermost hettangian or lowermost sinemurian, possibly corresponding to the j04 of partington et al. (1993) and the mfs indicated in the bucklandi zone by hesselbo & jenkyns (1998). sb 11 occurs in the bucklandi zone in a similar position to a major sequence boundary in the british basins (hesselbo & jenkyns 1998) and a maximum regression surface in the polish basin (feldman-olszewska 1997a); haq et al. (1988) indicated their ‘sb 202’ just below the hettangian–sinemurian boundary. mfs 11 in the turneri zone, sb 12 in the oxynotum zone and mfs 12 in the raricostatum zone all seem to be 1–2 ammonite zones out of phase relative to the haq et al. (1988) chart and surfaces identified in the british basins, the north sea and poland (fig. 29). their ages are based mainly on ostracod zonation, however, and 510 there are some uncertainties related to the correlation to ammonite zones at these levels. for instance, the p. reticulata subzone is correlated with the turneri zone by michelsen (1975), with the turneri–obtusum zones by sivhed (1980) and with the obtusum zone by park (1987). furthermore, the position of mfs 11 has been difficult to define in the distal and best-dated successions of homogeneous mudstones. these uncertainties may explain some of the difficulties in correlation of the sinemurian surfaces from the danish basin to skåne, poland and britain. in contrast, the lower pliensbachian – lower toarcian sb 13, mfs 13, sb 14, mfs 14, sb 15 and mfs 15 all occur in the same ammonite zones as the corresponding surfaces on the haq et al. (1988) chart and correlate quite well with the surfaces of partington et al. (1993), feldman-olszewska (1997a) and hesselbo & jenkyns (1998). three middle toarcian – lower aalenian sequences, fj 8–10, are preserved in the sorgenfrei–tornquist zone, corresponding to three sequences on the haq et al. (1988) chart. partington et al. (1993) indicated a maximum flooding surface in the uppermost toarcian, whereas feldman-olszewska (1997a) and hesselbo & jenkyns (1998) indicated a sequence boundary at this level. sb 19 corresponds precisely to the basinwards shift in facies that underhill & partington (1993, 1994) indicated at the top of the opalinum zone, coinciding with ‘sb 177’ of haq et al. (1988). the middle jurassic sequences in the danish basin are few compared to haq et al. (1988); they are poorly dated and their position in the scheme is uncertain. the upper oxfordian mfs 22 may correspond to ‘mfs 147’ and j54b, and sb 23 may correspond to ‘sb 146.5’. comparison of sea-level curves a relative sea-level curve has been constructed for the central parts of the danish basin including the sorgenfrei–tornquist zone based on the sequence stratigraphic framework (fig. 30). sequence boundaries and maximum flooding surfaces correspond to curve segments from inflection points to lows and tops, respectively, and the overall trend of the curve is determined by the areal extent of the sequences and the palaeowater depth indicated by facies. the lower jurassic sequences were partly removed by erosion in the southwestern parts of the basin owing to uplift, and middle jurassic sediments were not deposited there or over the ringkøbing–fyn high. the skagerrak–kattegat platform experienced relatively slow subsidence and parts of the upper pliensbachian were thinly developed and eroded, and upper aalenian – bajocian sed511 callovian bajocian pliensbachian hettangian norian toarcian aalenian bathonian oxfordian kimmeridgian sinemurian rhaetian sea-level rise poor age control a: hallam 1988 b: haq et al. 1988 c: surlyk 1990 d: this study e: hesselbo & jenkyns 1998 regional sea level east greenland sea level uk relative sea level danish basin eustatic sea level fig. 30. comparison of proposed eustatic and regional sea-level curves with the relative sea-level curve for the danish basin. curves a–c have been re-scaled according to the time-scale of gradstein et al. (1994). note the overall positive correlation for the lower jurassic; the middle jurassic part of the danish curve is poorly constrained. iments were probably not deposited (fig. 26). hence, the toarcian part of the sea-level curve is based primarily on data from the himmerland graben and the sorgenfrei–tornquist zone, whereas the middle jurassic part is based primarily on data from the sorgenfrei– tornquist zone. thus, strictly speaking, the whole curve is only valid for the sorgenfrei–tornquist zone. the late triassic – early jurassic part of the curve is, however, valid for most of the basin as subsidence was relatively uniform during this time interval. four sea-level curves are selected for detailed comparison (fig. 30). the curves of hallam (1988) and haq et al. (1988) are claimed to be essentially eustatic. the curve of hesselbo & jenkyns (1998) is based on the british lower jurassic, and the east greenland curve of surlyk (1990, 1991) represents the western margin of the major jurassic seaway of which the danish basin formed a part. it is generally accepted that a significant transgression occurred in northern europe shortly before the end of the triassic followed by a rapid regression more or less at the triassic–jurassic boundary (hallam 1992, 1997). the well-preserved succession in the danish basin, however, indicates that the late triassic transgression occurred in two major steps, one in the late norian and one in the latest rhaetian, separated by a significant fall in the late early rhaetian. in addition, the sea-level fall at the rhaetian–hettangian boundary seems to be of much smaller magnitude than that shown by hallam (1988) and haq et al. (1988), and suggested by sections in poland and germany, where significant fluvial incision is observed (bloos 1990; pieńkowski 1991; hallam 1992). a sea-level rise followed very shortly afterwards, in the planorbis zone. the danish curve indicates two maxima, one in the middle hettangian and one in the late hettangian, while the other curves only indicate one in the middle–late hettangian. all five curves indicate a fall close to the hettangian–sinemurian boundary. an extensive transgression began in the danish basin in the early sinemurian bucklandi zone (probably the upper part), as also recorded in the british curve. the sea-level rise is more marked on the hallam (1988) curve than on the haq et al. (1988) curve. in germany, two periods of rapid sea-level rise have been recognised in the early jurassic, in the planorbis subzone and the bucklandi subzone (bloos 1990), as also evident in the danish basin. the mid-sinemurian maximum and late sinemurian low of the danish curve is slightly out of phase with the hallam (1988) and haq et al. (1988) curves, which may reflect uncertainties in the biostratigraphic dating, as previously mentioned. the detailed british curve shows two general maxima and two general lows, of which the late sinemurian low may correspond to that recorded in the danish basin. a rise beginning in the latest sinemurian is shown by both the danish and british curves; surlyk (1990) shows a large rise at the sinemurian–pliensbachian boundary, whereas hallam (1988) and haq et al. (1988) indicate that the rise first started in the early pliensbachian. all the curves show a low in the midto late pliensbachian (at or close to the boundary of the davoei–margaritatus zones), and they all show a maximum in the uppermost pliensbachian (spinatum zone). the danish, british, and eustatic curves all agree on a rapid rise in the early toarcian, and a major fall is indicated in the early aalenian by the eustatic curves and the danish curve. the middle jurassic part of the danish curve is steep and smooth reflecting the thin and poorly dated sequences of proximal facies. the corresponding part of the other curves are quite different. the hallam (1988) and surlyk (1990) curves emphasise a bathonian rise, whereas the haq et al. (1988) curve shows a large fall. the late callovian – kimmeridgian part of the danish curve shows a rise, as in the other curves. the differences in the early jurassic sea-level curves are relatively small and may reflect different emphasis in the interpretations, uncertainties in the dating or local subsidence patterns in the danish basin. however, the general positive correlation with the other curves favours a eustatic cause for the long-term trend. furthermore, some of the short-term sea-level changes, for instance the maxima of the middle–late hettangian, the early pliensbachian, the late pliensbachian and the early toarcian, and the lows at the end of the early pliensbachian and the end of the late pliensbachian may also reflect eustatic changes, if the hallam (1988) and haq et al. (1988) curves reflect real eustatic changes. the curves may in part be biased by regional north european sea-level changes, caused by large scale intraplate stresses of the kind proposed by cloetingh (1988) rather than global changes (see discussion in miall 1994). in contrast, the middle jurassic part of the danish curve clearly reflects the tectonic influence involving regional uplift of the ringkøbing–fyn high, north-eastwards tilting of the basin and strong reduction in subsidence rate in the sorgenfrei–tornquist zone. similar and roughly contemporaneous tectonic influences have been recognised in other parts of the jurassic seaway (hallam & sellwood 1976; surlyk 1977, 1978; eynon 1981; underhill & partington 1993). approximately 150–170 m of paralic sediments were deposited in the 512 sorgenfrei–tornquist zone and a large proportion of the necessary accommodation space could be related to eustatic sea-level rise, if the curve of haq et al. (1988) is correct. correlation from the deep basin to the basin margin the rhaetian to early jurassic was a period of relative tectonic tranquillity and uniform subsidence. deposition was to a large extent controlled by regional sea-level changes that influenced most of the depositional basin including the basin margins in skåne and bornholm. a chronostratigraphic scheme of the danish basin and fennoscandian border zone is presented here (fig. 31) showing the rhaetian (in part) to lower aalenian sequence stratigraphic key surfaces, depositional environments, biozones and lithostratigraphy to facilitate comparison of the deeply-buried marine basinal successions of the danish basin with marginal successions in skåne and bornholm. the sequence stratigraphic interpretation of the hettangian–aalenian succession from bornholm is based on koppelhus & nielsen (1994), nielsen (1995), petersen & nielsen (1995) and surlyk et al. (1995), while detailed sequence stratigraphic interpretations of the succession in skåne are only documented for the sinemurian (erlström et al. 1999; frandsen & surlyk 2003, this volume). the preliminary correlation to the danish basin is thus based on descriptions of the rhaetian–sinemurian depositional facies (troedsson 1951; pieńkowski 1991; ahlberg 1994; ahlberg & arndorff 1994; erlström & guy-ohlson 1999), and lithostratigraphy (sivhed 1984; norling et al. 1993; ahlberg et al. 2003, this volume). the pronounced fluvial incision surface (sb 5) in the lower rhaetian of the danish basin is correlated with the base of the bjuv member in skåne and it is suggested that coal seams a and b were formed during the middle–upper rhaetian fourth-order transgressive events. the unconformable base of the hettangian fluvial sandstones (boserup beds) of the helsingborg member is correlated with sb 9. the two hettangian maximum flooding surfaces (mfs 9, mfs 10) are difficult to identify with confidence from the available descriptions but may be correlated with marine intervals at the planorbis/liasicus zone boundary and in the angulata zone (norling et al. 1993). braided fluvial sandstones are recognised in the lowermost sinemurian of the lower döshult formation in the örby section (erlström et al. 1999; locality 2), and the base of the sandstones which form the boundary to the paralic helsingborg member is here correlated with sb 11 in the bucklandi zone. the paralic flooding surface on top of the fluvial sandstones at örby is correlated with ts 11 which was a very significant and rapid marine flooding event in the bucklandi zone that influenced the entire basin. the next sequence boundary in the örby section which is overlain by a thin fluvial sandstone may correspond to a minor event of coastal progradation in the fjerritslev trough and at stenlille probably occurring in the semicostatum zone (fig. 31). the maximum flooding surface occurring in the döshult member, 5 m below the base of the pankarp member (frandsen & surlyk 2003, this volume) is correlated with mfs 11, a correlation that is supported by the ostracod data. the boundary between the döshult and pankarp members is interpreted as a sequence boundary by frandsen & surlyk (2003, this volume) and is correlated with sb 12, based on the available biostratigraphy. a regressive event in the late sinemurian is indicated by a sandstone layer with a coal bed in the pankarp member (sivhed 1984; norling et al. 1993); this event is not recognised in the deep basin. sequence stratigraphic correlation with the remaining part of the marine succession, the katslösa and rydebäck formations, is not yet possible. it is suggested, however, that the conglomeratic horizons indicating breaks in the marine rydebäck member (norling et al. 1993) are likely candidates for sequence boundaries. the hettangian–sinemurian sose bugt section of the arnager–sose fault block unconformably overlies upper triassic strata marking sb 1 of surlyk et al. (1995), which is correlated with sb 9 of this study. furthermore, the lower sinemurian sb 2 of surlyk et al. (1995) is correlated with sb 11, and the upper sinemurian sb 4 with sb 12 (fig. 31). the upper hettangian sb 10 of this study is proposed to separate the lower hettangian munkerup member from the upper hettangian – sinemurian sose bugt member; the boundary between the two members is not known, however, due to lack of exposure. the questionable sb 3 recognised by surlyk et al. (1995) at the base of some small channel units may correspond to the above-mentioned short-term event of coastal progradation in parts of the danish area (stenlille) and in the fjerritslev trough. mfs 10 and mfs 11 are correlated with thin shoreface sands representing marine incursions and mfs 12 is correlated with a 20 cm thick clay bed occurring just above sb 4 of surlyk et al. (1995). sb 13 is assumed to occur in the marine shoreface sandstones of the hasle formation and may correlate to an erosion surface in the type section of the hasle formation separating sandstones with two 513 514 bo un di ng s ur fa ce s li th os tr at ig ra ph ic u ni ts se qu en ce bo un da ry m ax im um m ar in e flo od in g su rf ac e t ra ns gr es si ve su rf ac e g a g as su m f or m at io n fia fib m em be rs in fii fj er ri ts le v fo rm at io n fiii fiv bj bj uv m em be r bo bo se ru p be ds h e h el si ng bo rg m em be r d ö d ös hu lt m em be r pa pa nk ar p m em be r k a k at sl ös a m em be r r y r yd eb äc k m em be r r ø r øn ne f or m at io n h a h as le f or m at io n so so rt ha t fo rm at io n ba ba gå f or m at io n sb m fs t s nannoceratopsis gracilis r. r ha et ic a po or r ec or ds c. b et zi . – c . c ra ss ire tic . d an is h ba si n fe nn os ca nd ia n bo rd er z on e sequence strat. key surfaces st en lil le w el ls v ed st ed -1 h al da ge r1 bø rg lu m -1 fl yv bj er g1 r øn ne g ra be n a rn ag er –s os e fa ul t bl oc k fj er ri ts le v t ro ug h h im m er la nd g ra be n sk ån e bo rn ho lm toarcian pliensbachianaal. sinemurian hettangian rhaet.ages o pa lin um a al en si s ps eu do ra di os a d is pa ns um t ho ua re nc e v ar ia bi lis bi fr on s fa lc ife ru m t en ui co st at um sp in at um m ar ga ri ta tu s d av oe i ib ex ja m es on i r ar ic os ta tu m o xy no tu m o bt us um t ur ne ri se m ic os ta tu m bu ck la nd i a ng ul at a li as ic us pl an or bi s m ar sh i m ur ch is on ae sequences h a 1 f 10 fj 9 fj 8 fj 7 fj 6 fj 5 fj 4 fj 3 fj 2 fj 1 g a 1 b ababa m . s em ita bu la tu m p. n as ut a l. s pi no sa m . se m ita bu la tu m d . p ris cu m l. v ar ia bi le ‘lo w er su bz on e’ ‘m id dl e su bz on e’ p. r et ic . ‘u pp er su bz on e’ g . a po st ole sc ui – k. (k .) fo ve ol at a ‘u pp er su bz on e’ ‘lo w er su bz on e’ o . d an ic a o . a sp in at a o . a de nt ic ul at a – n . ( n .) sim pl ex ri cc ii. –p ol yp od ii. rh ae ti. –l im bo . m io sp or e zo ne s d in of la ge lla te cy st z on es po ul se n 19 96 o st ra co d zo ne s c hr on oz on es sb 1 9 m fs 1 8 m fs 1 7 m fs 1 6 m fs 1 5 m fs 1 4 m fs 1 3 m fs 1 2 m fs 1 1 m fs 1 0 m fs 9 m fs 7 t s 9 t s 10 t s 11 t s 12 t s 13 t s 14 t s 15 sb 1 8 sb 1 7 sb 1 6 sb 1 5 sb 1 4 sb 1 3 sb 1 2 sb 1 1 sb 1 0 s b 9 sb 5 fl uv ia l s an ds to ne s es tu ar in e/ la go on al h et er ol ith s, m ud st on es , c oa l b ed s sh or ef ac e sa nd st on es /s ilt st on es es tu ar in e ch an ne l s an ds to ne s o ffs ho re m ud st on es d ep os iti on al e nv ir on m en t o ffs ho re – lo w er s ho re fa ce he te ro lit hs , m ud dy s an ds to ne s la cu st ri ne m ud st on es a nd co al b ed s bo b ed s d ö m b h e m b bj m b r ø fm r ø fm h a fm h a fm so f m ba f m r y m b k a m b pa m b fiv m b fiii m b fii m b fib m b g a fm fia m b fiv m b h a fm fiii m b fii m b fib m b fia m b g a fm fiv m b fiii m b fii m b fib m b fia m b fiii m b fii m b fib m b fia m b g a fm g a fm pe rin op ol le ni te s el at oi de s pi nu sp ol le ni te s– tr ac hy sp or ite s ce re br op ol le ni te s m ac ro ve rr uc os us sp he rip ol le ni te s– le pt ol ep id ite s different diagenetic evolutionary trends (larsen & friis 1991; nielsen 1995); biostratigraphic evidence from ammonites seems to confirm this suggestion (donovan & surlyk 2003, this volume). the base of the fluvial sand in the lower sorthat formation (michelsen et al. 2003, this volume) is correlated with sb 14, and mfs 14 is correlated with a c. 2.1 m thick lagoonal unit with carbonaceous clay containing abundant dinoflagellates and a marine-influenced coal seam (batten et al. 1994; koppelhus & nielsen 1994; nielsen 1995; petersen & nielsen 1995). sb 15 occurs at the base of a 19 m thick fluvial–estuarine channel unit and mfs 15 is correlated with hummocky cross-stratified sands occurring in a lagoonal unit, based on the dinoflagellate biostratigraphy (koppelhus & nielsen 1994). the base of a thin fluvial sand possibly correlates with sb 16. an upper toarcian hiatus is suggested from the available biostratigraphy. the 4–5 m thick unit of transgressive shoreface sandstones that erosionally overlies lacustrine clays is correlated with the early aalenian mfs 18, and the base of the overlying coarse-grained pebbly alluvial sands is correlated with sb 19, which marks the significant change in deposition to continental middle jurassic. synthesis of upper triassic – jurassic basin evolution basin formation and structural development late carboniferous continental deposition preceded the rifting of the basin (michelsen & nielsen 1991, 1993), which was presumably related to late carboniferous – early permian plate reorganisations (ziegler 1982; vejbæk 1989, 1990). regional crustal stretching and dextral strike-slip movements along the tornquist zone caused formation of extrusive and intrusive volcanic rocks, numerous tilted fault blocks composed of lower palaeozoic rocks and deposition of thick lower permian syn-rift prisms in the danish basin, the fennoscandian border zone and the oslo graben (vejbæk 1985, 1989, 1990; liboriussen et al. 1987; norling & bergström 1987; ro et al. 1990; michelsen & nielsen 1991, 1993; mogensen 1994, 1996). the thick wedges in the hangingwall blocks comprise alluvial conglomerates and lacustrine mudstones with volcaniclastics eroded from footwall crests (michelsen & nielsen 1991). continued erosion caused deep truncation of the crests and formation of the regional mid-permian unconformity that forms the base for the post-rift succession. possible phase transformations in the deep crust and post-rift cooling caused rapid regional subsidence in late permian – early triassic times (vejbæk 1989, 1990). the danish basin and parts of the sorgenfrei–tornquist zone were transgressed in the late permian, and thick carbonate and salt deposits accumulated in the centre of the basin, whereas shallow marine and continental clastics were deposited along the basin margin in the sorgenfrei–tornquist zone. during early–middle triassic times, several thousands of metres of dominantly continental strata were deposited in a hot and arid climate (bertelsen 1980). the triassic transtensional strike-slip movements in the sorgenfrei–tornquist zone continued into the early jurassic causing subsidence of the himmerland graben, the fjerritslev trough, the rønne graben and parts of skåne (rolle et al. 1979; gravesen et al. 1982; vejbæk 1985, 1990; liboriussen et al. 1987; norling & bergström 1987; erlström et al. 1997). minor down-to-basin fault displacements occurred locally along the southern basin margin. the thermal subsidence continued into late triassic – early jurassic times, and most of the basin experienced relatively uniform subsidence until the early aalenian uplift completely changed the configuration of the basin. late triassic marine flooding a gradual change from arid towards more humid conditions took place in the late triassic, partly due to formation of a large epicontinental sea (bertelsen 1978, 1980; parrish et al. 1982; ziegler 1982; hallam 1984, 1985; batten et al. 1994; scotese 1994). carnian deposition of lacustrine and sabkha mudstones was terminated by the early norian marine transgression that probably came from the south. the transgression led to deposition of oolitic limestones succeeded by marlstones and fossiliferous claystones of the vinding 515 facing page: fig. 31. chronostratigraphic scheme for the danish basin and the fennoscandian border zone showing the rhaetian (part) to lower aalenian sequence stratigraphic key surfaces, depositional environments, biozones and lithostratigraphy. the scheme compares the basinal successions of the deep danish basin with successions deposited along the basin margin in skåne and on bornholm during the late triassic – early jurassic period of relative tectonic tranquillity, widespread subsidence and maximum control by eustatic or regional sealevel changes. formation, 40–100 m thick (bertelsen 1978, 1980; l.h. nielsen, l. hamberg and e.b. koppelhus in: nielsen 1995). at its maximum late norian extent, the shallow sea covered most of the danish basin, the ringkøbing– fyn high and the north german basin. at the same time, fluvial arkosic sands and lacustrine muds were deposited in the sorgenfrei–tornquist zone, skåne, rønne graben and along the northern basin margin. these deposits are included in the skagerrak and kågeröd formations (bertelsen 1980; gravesen et al. 1982; sivhed 1984; ahlberg 1994; nielsen 1995). after the maximum transgression, a phased regression followed, and shoreface and fluvial sands of the lower gassum formation were deposited in stepwise more basinwards positions intercalated with clays of the upper vinding formation in the basin centre. in the fjerritslev trough, alternating units of arkosic fluvial sand and lacustrine mud of the skagerrak formation were deposited, reflecting repeated base-level changes controlled by the short-term sea-level fluctuations. the regression culminated in the early rhaetian with the formation of an extensive, fluvially-incised sequence boundary (sb 5; figs 26, 31). at this time, the ringkøbing–fyn high was exposed to erosion and fine-grained sand was shed to the basin from the high. non-deposition or erosion probably occurred on the skagerrak–kattegat platform, in skåne and on bornholm, east of the rønne graben. when sea-level slowly rose again, fluvial–estuarine deposits, up to 30 m thick, were deposited above the sequence boundary before widespread flooding occurred. the transgression continued but was punctuated by two short-term, forced regressions that led to deposition of widespread shoreface sand sheets encased in transgressive offshore mud. the transgression accelerated and reached its maximum in the latest rhaetian, when the danish basin, the sorgenfrei–tornquist zone, the skagerrak–kattegat platform, the north german basin, and the ringkøbing–fyn high were covered by the sea (mfs 7, figs 26, 31). the aggrading floodplain succession in skåne with the rhaetian coal seams a and b shows weak marine influence (vallåkra and bjuv members; troedsson 1948, 1951; lund 1977; sivhed 1984; norling et al. 1993; ahlberg & arndorff 1994) and was probably deposited during this transgressive phase (fig. 31). the increasing amount of macroplant fossils, rootlets, coal seams and more mature sandstones indicates that the climate became more humid during the late rhaetian transgressive phase. latest rhaetian – earliest hettangian regression an overall sea-level fall commenced after the maximum transgression, and two phases of coastal progradation, at the rhaetian–hettangian boundary and in the earliest hettangian, caused deposition of two thin regressive shoreface sand sheets that constitute the upper part of the gassum formation over much of the danish basin. the regression culminated with coastal progradation far into the basin accompanied by fluvial erosion and incision in the himmerland graben, the fjerritslev trough and along the southern basin margin, where parts of the ringkøbing–fyn high were exposed and supplied sand to the basin (sb 9; fig. 31). at the same time, rhaetian or older rocks were eroded in skåne, poland, north germany, england and on bornholm; the resultant widespread sequence boundary marking a late rhaetian hiatus is onlapped by transgressive lower jurassic deposits. thus, the relatively conformable upper triassic – lowermost jurassic succession in the danish basin reflects a more rapidly fluctuating sea level than typically recognised (hallam 1988, 1997) with maxima in late norian and latest rhaetian times and lows in the early rhaetian and the earliest hettangian. early jurassic basin expansion a subtropical to warm-temperate and humid climate characterised the jurassic period, and large quantities of clay were supplied to the basin from the weathering of palaeozoic shales, granitic basement of the baltic shield and a possible carboniferous regolith (ziegler 1982; pedersen 1983; schmidt 1985; norling & bergström 1987; nielsen & koppelhus 1991; surlyk et al. 1995). the basin expanded north-eastwards against the baltic shield during early jurassic time owing to the combined effects of eustatic sea-level rise, regional subsidence and local block-faulting caused by transtensional strike-slip movements in the sorgenfrei–tornquist zone (rolle et al. 1979; gravesen et al. 1982; norling & bergström 1987; nielsen 1995; surlyk et al. 1995; mogensen 1996). hettangian – early pliensbachian transgression in the danish and north german basins, the transgression commenced in the early hettangian planorbis 516 zone. fully marine mudstones belonging to the f-ia unit of the fjerritslev formation overlie the sandy gassum formation in most of the danish basin (fig. 31). the mudstones have a low content of ostracods, foraminifera and infaunal bivalves, but the high content of landderived organic matter suggests a relatively high rate of deposition. in the north-eastern part of the sorgenfrei–tornquist zone, aggrading parasequences of fluvial and shoreface sands with subordinate offshore muds were deposited, while lagoonal parasequences were formed on the skagerrak–kattegat platform. the transgression peaked in the early and late hettangian, interrupted by a short-term regression in the middle hettangian (sb 10; fig. 31). the depositional area also expanded in the southeasternmost part of the basin on bornholm. in the rønne graben, lacustrine conditions in the early hettangian were followed by deposition of dominantly coastal plain deposits, up to 200 m thick. onshore bornholm, lacustrine mudstones overstepped small fault blocks (gravesen et al. 1982). the basin expansion is also evident in skåne, poland and germany, where basal hettangian fluvial deposits are overlain by thick coastal plain, deltaic and shallow marine deposits, locally up to 400 m thick (troedsson 1951; norling 1972; lund 1977; sivhed 1984; pieńkowski 1991; ahlberg 1994; ahlberg & arndorff 1994; feldman-olszewska 1997a; ahlberg et al. 2003, this volume). the large increase in accommodation space governing deposition of transgressive paralic deposits along the basin margin was interrupted briefly by a fall in sea level, soon after the hettangian–sinemurian boundary. this resulted in fluvial incision on the skagerrak–kattegat platform and on bornholm and deposition of braided fluvial sandstones in skåne (basal döshult member), while regressive shoreface sand was deposited in the fjerritslev trough (sb 11; fig. 31). farther basinward, heteroliths and silty mudstones were deposited above the conformable part of the sequence boundary. a rapid sea-level rise followed in the earliest sinemurian (upper part of the bucklandi zone) and transgressive marine muds of the f-ib unit finally overstepped fluvial and marine sands of the gassum formation in the sorgenfrei–tornquist zone and on the skagerrak–kattegat platform. in the basin, up to 150 m of uniform mudstones were deposited, showing a marked thinning towards the north-eastern basin margin. the corresponding deposits in the rønne graben consist of c. 130 m of aggrading lagoonal parasequences with a marine shoreface unit marking the culmination of the transgression (mfs 11; fig. 31; nielsen 1995). on the shallow arnager–sose fault block, the culmination is recorded by a thin shoreface sandstone. in skåne, fluvial, estuarine and shallow marine sands and muds of the döshult member were deposited followed by marine clays of the pankarp member (norling 1972; sivhed 1980, 1984; norling et al. 1993; erlström et al. 1999; frandsen & surlyk 2003, this volume). in the late sinemurian, a minor sea-level fall caused a slight basinwards progradation of coastal parasequences on the skagerrak–kattegat platform and in parts of the sorgenfrei–tornquist zone. after this minor excursion, the overall early jurassic sea-level rise continued, and reached a maximum in the latest sinemurian, possibly close to the oxynotum–raricostatum zone boundary. in the centre of the danish basin, the diversity and abundance of the ostracod fauna decreased and infaunal bivalves and some of the epifaunal bivalves disappeared due to reduced oxygenation (pedersen 1986; michelsen 1989b). in the rønne graben, lagoonal and lacustrine parasequences with minor fluvial sand continued to aggrade (amounting to c. 200 m) until sea-level rise outpaced the rate of paralic sedimentation (nielsen 1995). restricted offshore mud was deposited on the southern part of the rønne–hasle fault block and on the arnager–sose fault block during the maximum phase of transgression (nielsen 1995; surlyk et al. 1995). a gradual decrease in the rate of sea-level rise in the early pliensbachian jamesoni zone caused the formation of thick, aggrading to weakly forestepping shallow marine parasequences on the arnager–sose and rønne– hasle fault blocks (the lower hasle formation), and on the skagerrak–kattegat platform, where a distinct basinwards progradation of shoreface sandstones is seen (fig. 25). the coastal sandstones capped by a coal bed of the upper pankarp member in skåne (sivhed 1984; frandsen & surlyk 2003, this volume) seem to represent a local regression based on the available biostratigraphic data. the regression culminated in the middle early pliensbachian (early ibex zone); in the danish basin, deposition changed from fine-grained mud (f-ib unit) to silty and sandy heteroliths (f-iia beds, sb 13; fig. 31). in skåne, a distal sequence boundary was developed within marine mudstones separating the pankarp member from the overlying döshult member (frandsen & surlyk 2003, this volume). on the skagerrak–kattegat platform and on bornholm, deposition of shallow marine shoreface sand ceased for some time due to submarine or subaerial erosion and bypass. when the sea level started to rise again, deposition of fine-grained mud resumed in the danish basin (lower part of f-iib beds), while shoreface sand again was deposited on bornholm 517 (upper hasle formation). backstepping parasequences of marine sand were succeeded by transgressive mud on the skagerrak–kattegat platform. peak transgression was reached in the late early pliensbachian davoei zone. thereafter, the rate of sea-level rise decreased and a coarsening-upwards succession of mud and finegrained heteroliths was deposited in the danish basin (middle part of f-iib), while deposition of regressive marine sands occurred on bornholm. late pliensbachian – early aalenian sea-level fluctuations significant erosion took place on the skagerrak–kattegat platform during a sea-level fall in the early late pliensbachian margaritatus zone (sb 14; figs 25, 31). basinwards, in the fjerritslev trough and the danish basin, deposition changed to silty and sandy mud and finegrained sand, showing very marked thinning over salt structures possibly reflecting shallow water depths (upper part of f-iib and f-iic beds). at the basin margin, deposition of regressive shallow marine sands continued in skåne (rydebäck member), while marine deposition ceased in the rønne graben (nielsen 1995). the ensuing sea-level rise, which initiated the second early jurassic transgression, commenced in the margaritatus zone and reached a peak in the late late pliensbachian (early spinatum zone). marine silty mud was deposited in the danish basin, while marine sand with bivalves was deposited in the fjerritslev trough. deposits from this period are absent on the skagerrak–kattegat platform due to bypass or later erosion. on bornholm, the sealevel rise created accommodation for the accumulation of 100–140 m of coarse-grained fluvial sands, floodplain deposits, peats and lagoon-fill deposits referred to the new sorthat formation (fig. 31; koppelhus & nielsen 1994; nielsen 1995; petersen & nielsen 1995; michelsen et al. 2003, this volume). the following sea-level fall culminated in the late spinatum zone with the formation of a widespread regressive surface of marine erosion and progradation of a sandy shoreface into the central parts of the fjerritslev trough (sb 15; fig. 31). the sea-level fall caused erosion of lower pliensbachian strata on the skagerrak–kattegat platform, and incision on bornholm. the ensuing sea-level rise caused marine flooding over the entire basin at the pliensbachian–toarcian boundary, including the sorgenfrei–tornquist zone and the skagerrak–kattegat platform, and deposition of transgressive marine mud commenced. lagoonal sand and mud and restricted marine shoreface sand were deposited on bornholm (upper sorthat formation; koppelhus & nielsen 1994). the transgression reached its maximum in the early toarcian falciferum zone (mfs 15). due to oxygen-poor conditions, the ostracod fauna disappeared and an increasing amount of amorphous marine matter is preserved making the deposits a favourable source rock for hydrocarbons (thomsen et al. 1987; michelsen 1989b; dybkjær 1991). during the remainder of the early jurassic and in the early aalenian opalinum zone, a succession of up to 150 m of marine mudstones with three shoreface sandstones was deposited in the sorgenfrei–tornquist zone; the shoreface sandstones overlie regressive surfaces of marine erosion and accumulated during sea-level falls. on the skagerrak–kattegat platform, where accommodation was much less, thin peat-bearing lagoonal successions were deposited on the subaerial erosion surfaces during ensuing transgressions. the basin gradually shrank and became isolated from the north sea basin due to initial uplift, and a poor circulation pattern and possibly brackish conditions at the margin caused impoverishment of the fauna. in skåne, regressive, shallow marine sand of the rydebäck member was deposited, while erosion probably prevailed during much of late early – late toarcian time on bornholm (nielsen 1995). late early – middle jurassic uplift and erosion the ringkøbing–fyn high and most of the danish basin were uplifted in late early jurassic – early middle jurassic times, and the triassic – lower jurassic successions were eroded on the highest parts of the ringkøbing–fyn high. the lower jurassic was deeply eroded in the uplifted area north of the high, whereas erosion did not reach such deep levels closer to the sorgenfrei–tornquist zone. in the fault-bounded sorgenfrei–tornquist zone, where subsidence still occurred but at a much lower rate than before, the change in basin configuration resulted in a shift from deposition of homogeneous offshore muds to shallow marine sands. the sandstones overlie a forced regressive surface of marine erosion dated to the top of the lower aalenian opalinum zone, and are thus contemporaneous with the greatest basin shift in facies recorded from the lower–middle jurassic in the north sea basins. in skåne, uplift was accompanied by faulting, erosion and volcanism. hence, during the rest of the aalenian, the bajocian and the early bathonian, deposition was more or less confined to the narrow zone bounded by the fjerritslev and børglum 518 faults and their south-eastwards continuation in kattegat, øresund and skåne. material was supplied from both the uplifted areas to the west and south-west, and from the baltic shield. from the fjerritslev trough in the north-west to skåne in the south-east, shallow marine, paralic and fluvial sediments were deposited, making up the haldager sand formation, the vilhelmsfält formation and the mariedal formation. in the rønne graben, lacustrine muds, peats and alluvial sands of the bagå formation were deposited. the general scarcity of marine fossils and pronounced reworking of palynomorphs hinder precise dating of the deposits, and large parts of the aalenian–callovian time interval are probably not represented in the sorgenfrei–tornquist zone because of insufficient accommodation space and erosion during formation of sequence boundaries and ravinement surfaces. late middle – late jurassic basin expansion the area of subsidence gradually expanded in bathonian–volgian times with deposition of bathonian(?) braided fluvial sands on the skagerrak–kattegat platform and south-west of the fjerritslev trough in the himmerland graben. a marine transgression close to the callovian–oxfordian boundary influenced most of the basin and accommodation space was also created in the former by-pass zone of the southern part of the basin and on the skagerrak–kattegat platform, where fluvial sands were now deposited. during the oxfordian, the sedimentation area was further enlarged and a northeastwards thickening wedge of transgressive, fossiliferous marine sand and mud was deposited above lagoonal deposits on the skagerrak–kattegat platform. lagoonal deposition apparently still dominated to the south-west (flyvbjerg formation), where the structural high was still present, albeit with much reduced relief. a latest oxfordian sea-level fall resulted in coastal progradation on the skagerrak–kattegat platform and in the fjerritslev trough; fluvial and shallow marine sands were deposited, and a south-west prograding wedge was formed. extensive marine flooding occurred in the kimmeridgian, and sedimentation of marine mud (børglum formation) characterised the whole area, although the marked thinning towards the south-west emphasises the reduced accommodation here. during volgian–ryazanian times, the depositional environment was dominantly a shallow shelf with three–four major phases of coastal progradation (sequences fr 1, fr 2, fr 3; fig. 26). coastal and deltaic sandy deposits (frederikshavn formation) accumulated on the skagerrak–kattegat platform and in parts of the sorgenfrei–tornquist zone, while marine muds were deposited over much of the basin. the occurrence of sandy beds in the south-westernmost parts of the basin indicates that during low sea level, the ringkøbing–fyn high still supplied some sand. ammonites indicating the middle volgian albani zone (birkelund & pedersen 1980) are found in sequence fr 1, and dinocysts indicating the jurassic–cretaceous boundary (davey 1982) occur in sequence fr 2 close to the maximum flooding surface. on the bornholm high to the east, fluvial deposition followed by paralic deposition was resumed above a callovian–kimmeridgian unconformity due to the late volgian – ryazanian sea-level rise. conclusions the depositional area covering the danish basin and the fennoscandian border zone underwent great changes in late triassic – jurassic times in terms of tectonics, areal extent, sea level and climate, and these factors governed important differences in the depositional style of the upper triassic – jurassic sequences. a shallow, low-gradient marine embayment without a shelf-slope break was established in the danish basin and parts of the fennoscandian border zone in late triassic time. due to the gentle basin physiography, fourth-order sea-level changes exerted a strong control on the distribution and architecture of the depositional facies. extensive coastal progradation occurred during sea-level fall, and regressive shoreface sandstones were deposited over large areas, resting on regressive surfaces of marine erosion. subsequent sea-level rises caused widespread flooding and the formation of ravinement surfaces overlain by marine mudstones. the changes in stacking pattern from forestepping to backstepping and vice versa occur at a pronounced, third-order fluvial incision surface with thick estuarine valley-fill deposits, and at a widespread marine maximum flooding surface, respectively. the extensive fourth-order shoreface sandstone sheets and associated paralic deposits were developed because the shallow basin was influenced by rapid, low-amplitude sea-level fluctuations, a high, continuous sediment influx from the basin margins on three sides, and a high, longshore sediment distribution rate. fluvial transport of sand into the basin was effective, probably because of relatively poorly developed vegetation. a deeper and fully marine, low-gradient shelf was established in early jurassic time due to steady subsidence 519 and overall eustatic/regional sea-level rise. the ringkøbing–fyn high was permanently flooded, and the basin was primarily sourced with mud at a high continuous rate from landmasses to the east and north exposed to extensive weathering in the humid, warm-temperate to subtropical climate. at this time, vegetation was widespread and well-established. sedimentation kept pace with the steady formation of new accommodation space, and biozones and facies packages are thickly developed; omission surfaces and condensed sections are poorly developed. storm-generated seawards flowing currents carried mud in suspension into the basin, where it was deposited below average storm wave base. local topography was smoothed out by draping mud, forming widespread uniform and thick packages of mudstones with a conformable stratal pattern. minor sea-level fluctuations left only insignificant traces compared to the upper triassic because of the relatively deep depositional basin. only when larger eustatic falls exceeded subsidence and wave-base was significantly lowered were coarse-grained mudstones and fine-grained heteroliths deposited above regressive, marine surfaces; these deposits form thin lowstand systems tracts overlying conformable sequence boundaries. coastal sand was mainly trapped at the basin margin and subaerial exposure was very limited within the study area. the sea-level curve from the danish basin parallels the trend of other early jurassic sea-level curves and shows great similarities in detail, testifying to the eustatic or large-scale regional influence of sea level on the depositional pattern. the regime of tectonic tranquillity that characterised most of the basin during the late triassic – early jurassic was replaced by one of uplift of most of the basin in middle jurassic time. it began with a reduction of the depositional area and the development of a restricted marine environment in the late toarcian, and culminated with the formation of a regional intra-aalenian unconformity, the most significant unconformity of the upper triassic – jurassic succession. deposition in late aalenian – bajocian times became confined to the faultbounded sorgenfrei–tornquist zone, where subsidence still occurred, albeit at a much reduced rate. the preserved sequences consist primarily of fluvial–estuarine and lagoonal deposits of late lowstand to early transgressive systems tracts, while late transgressive and highstand systems tracts are almost absent, due to cannibalisation during sea-level falls owing to the very slow subsidence rate. during the late jurassic, subsidence accelerated and the depositional area gradually increased in size and depth, although the ringkøbing– fyn high functioned as a low-relief paralic hinterland that still supplied small amounts of sand to the basin until the ryazanian. the uplift of the ringkøbing–fyn high and its flanks including most of the danish basin occurred contemporaneously with uplift in the central north sea. the intra-aalenian unconformity shows progressively deeper truncation against the high and progressively younger strata onlap the unconformity in the same direction, similar to the pattern shown by the ‘mid-cimmerian unconformity’ in the north sea against the dome centre. the similarity in age and geometric configuration points to a common causal mechanism for the formation of these unconformities. the subcrop and onlap contour patterns of the unconformity parallel the ringkøbing–fyn high and are not compatible with the postulated semi-circular subcrop stratigraphy of the north sea unconformity that has been used as an argument for a mantle plume causing domal uplift. the uplifted area rather takes the form of a broad irregular shaped arch. the contour pattern in the danish basin, the occurrence of several volcanic centres in the north sea and the presence of middle jurassic basalts in southern sweden indicate that an interpretation of the regional uplift should consider a more complex plume pattern, in addition to the response of the ringkøbing–fyn high and the role of the deep, fundamental sorgenfrei–tornquist fracture zone. acknowledgements the paper is based on a ph.d. study at the university of copenhagen supervised by finn surlyk and olaf michelsen. the supervisors, colleagues at geus (jan andsbjerg, torben bidstrup, karen dybkjær, peter japsen, peter n. johannessen, jon r. ineson, anders mathiesen, niels e. poulsen, ole v. vejbæk), lars hamberg (dong) and eva b. koppelhus are thanked for stimulating discussions. the study was financially supported by geus and the danish research academy. the reviews of stephen hesselbo, jon r. ineson and gunver k. pedersen are highly appreciated. 520 references ahlberg, a. 1994: facies analysis of alluvial and deltaic environments in the hettangian of nw skåne, southern sweden, 18 pp. in: ahlberg, a.: deposition and diagenesis of the rhaetian– hettangian succession (triassic–jurassic) in southern sweden: a synthesis. lund publications in geology 123. ahlberg, a. & arndorff, l. 1994: pedogenesis and sedimentology of alluvial upper triassic (middle rhaetian) strata of the bjuv member (höganäs formation), southern sweden, 17 pp. in: ahlberg, a.: deposition and diagenesis of the rhaetian– hettangian succession (triassic–jurassic) in southern sweden: a synthesis. lund publications in geology 123. ahlberg, a., sivhed, u. & erlström, m. 2003: the jurassic of skåne, southern sweden. in: ineson, j.r. & surlyk, f. 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(eds): tectonic evolution of the north sea rifts, 1–36. oxford: clarendon press. manuscript received 5 february 1999; revision accepted 24 june 1999. oa 4 rp rl ce rl rl ce rl rp 9 oa rp rr rp rr pt rl ce rp8 4 mfs 1 mfs 9 skagerrak–kattegat platform ? skagerrak formation sorgenfrei–tornquist zone 19 km 25 km 37 km 5 km 41 km 30 km 20 km børglum fault sb 1–7 sb 9 mfs 7 sb 8 ts 8 ts 9 mfs 9 sb 9 ts 9 mfs 7 sb 8 sb 6 sb 4 mfs 1 mfs 8 sb 7 ts 7 mfs 6 mfs 5 ts 5 sb 5 mfs 4 sb 3 mfs 3 ts 4 oddesund formation ts 1 pt pt ts 7 himmerland graben sæby-1 gr vedsted-1 sp mejrup-1 gr rødding-1 sp hyllebjerg-1 gr farsø-1 grvemb-1 sp flyvbjerg-1 sp nesw dp ce rl oa rl rr rl/rp pt 3 10 oa dp pt sb 2 lst hst tst hst tst lst hst tst rl 9 fj 1 ga1 vi 1 biostratigraphy spore/pollen zonation pt pinuspollenites–trachysporites zone rp ricciisporites–polypodiisporites zone rl rhaetipollis–limbosporites zone ce corollina–enzonalasporites zone ostracod zonation oa ogmoconchella aspinata zone dinoflagellate zonation dp dapcodinium priscum zone rr rhaetogonyaulax rhaetica zone depositional environments fluvial estuarine lacustrine lagoonal shoreface offshore 50 m vemb-1 vedsted-1 sæby-1 50 km sorgenfrei–tornquist zone børglum fault haldager fault fjerritslev fault danish basin skagerrak–kattegat platform flyvbjerg-1 mejrup-1 rødding-1 farsø-1 hyllebjerg-1 ringkøbing–fyn high systems tracts lst lowstand tst transgressive hst highstand bounding surfaces sb sequence boundary mfs maximum marine flooding surface ts transgressive surface core (core numbers) vi 1 sequence fig. 19. sw–ne well-log panel across the danish basin, the sorgenfrei–tornquist zone and the skagerrak–kattegat platform showing two third-order sequences, the norian – lower rhaetian vi 1 bounded by ts 1 and sb 5, and the rhaetian – lowermost hettangian ga 1 bounded by sb 5 and sb 9. the lower part of the hettangian third-order fj 1 sequence is shown in the upper part of the panel. the tst of the vi 1 sequence consists of simple backstepping parasequences, while the hst shows a more complex development with three fourth-order sequences bounded by sb 2 through sb 5. similarly, the ga 1 sequence consists of four fourth-order sequences bounded by sb 5 through sb 9. the third-order surfaces marked in bold (i.e. ts 1, mfs 1, sb 5, mfs 7, sb 9, ts 9, mfs 9) are defined by changes in the stacking pattern of the fourth-order sequences; this is most clearly seen in distal successions such as in the vemb-1, mejrup-1 and rødding-1 wells. modified from l.h. nielsen, l. hamberg and e.b. koppelhus (in: nielsen 1995) and hamberg & nielsen (2000). mejrup-1 gr hyllebjerg-1 gr vedsted-1 sp flyvbjerg-1 sp frederikshavn-2 sp børglum-1 sp børglum fault vedsted-1 flyvbjerg-1 børglum-1 50 km sorgenfrei–tornquist zone børglum fault haldager fault fjerritslev fault danish basin skagerrak–kattegat platform mejrup-1 frederikshavn-2 hyllebjerg-1 ringkøbing–fyn high 100 m depositional environments fluvial estuarine lacustrine lagoonal shoreface offshore fjerritslev trough himmerland graben skagerrak–kattegat platform 65 km 41 km 30 km 15 km 42 km 5 6 7 8 ts 7 ts 9 mfs 7 sb 12 sb 11 sb 8 ts 11 mfs 9 mfs 11 mfs 12 ts 12 sb 9 mfs 7 sb 11 sb 10 ts 10 sb 9 sb 8 ts 9 mfs 8 ts 11 ts 7 mfs 9 mfs 10 mfs 11 mfs 12 sb 12 sw ne 17 8 6 7 8 7 lst lst tst lst hst tst hst/frst tst hst tst frst lst tst lst lst frst hst tst tst hst lst hst 18 19 20 21 ga 1 fj 4 fj 3 fj 2 fj 1 systems tracts lst lowstand tst transgressive hst highstand frst forced regressive bounding surfaces sb sequence boundary mfs maximum marine flooding surface ts transgressive surface core fj 1 sequence fig. 20. sw–ne well-log panel across the danish basin to the skagerrak– kattegat platform showing the upper part of the rhaetian – lowermost hettangian sequence ga 1 and the hettangian–sinemurian sequences fj 1, fj 2, fj 3 and the lower part of fj 4. note the thickly developed hst of sequence fj 1 in børglum-1 and flyvbjerg-1 on the basinwards side of the børglum fault. the interbedding of estuarine sandstones with marine mudstones and sandstones indicates repeated fluctuations of sea level during the formation of the hst and suggests the presence of several fourth-order sequences at this level. skive-1 ts 10 sb 10 m l k j i f e d c a iia iib iic pll plu jl su sl h fiii m em be r fii m em be r fia m em be r fib m em be r fiv m em be r fj er ri ts le v fo rm at io n sonicgrsonicgrsonicgrsonicgrsonicgr oddesund-1 rødding-1 skive-1 kvols-1 hyllebjerg-1 sb 19–22 mfs 11 mfs 12 sb 12 sb 14 mfs 13 sb 13 sb 11 ts 11 sb 11 ts 11 sb 10 ts 10 mfs 11 mfs 10 sb 12 ts 12 mfs 12 sb 13 ts 13 mfs 13 sb 14 ts 14 mfs 14 sb 15 ts 15 ts 16 sb 16 mfs 15 mfs 16 plu pll sb 19–21 plu plu pll sl plu pl pll su su su sl–h sw ne fj 3 fj 4 fj 5 fj 6 fj 7 fj 8 fj 1 fj 2 lst lst tst tst hst fj 2 fj 3 fj 5 fj 4 fj 6 fj 1 18 km 12 km 19 km 32 km hst hst tst hst tst lst lst hst tst hst lst lst tst hst lst tst sb 10 a sb 11 sb 12 g sb 13 sb 14 sb 15 h d c b a e g b g h 100 m depositional environment offshore stratigraphy lithostratigraphic boundary strata boundary (a–m) after michelsen (1989b) single dated samples top of chronostratigraphic unit, based on a series of ostracod analysis top of ogmoconchella aspinata zone jl lower jurassic pl pliensbachian plu upper pliensbachian pll lower pliensbachian su upper sinemurian sl lower sinemurian h hettangian systems tracts lst lowstand tst transgressive hst highstand bounding surfaces sb sequence boundary mfs maximum marine flooding surface ts transgressive surface fj 1 sequence 50 km sorgenfrei–tornquist zone børglum fault haldager fault fjerritslev fault danish basin skagerrak–kattegat platform oddesund-1 rødding-1 hyllebjerg-1 kvols-1 ringkøbing – fyn high fig. 21. well-log panel showing the basinal part of the sequences fj 1 to fj 8 from hyllebjerg-1 (himmerland graben) towards more basinal well sections. an increasing amount of erosion is seen towards the south-west below the ‘base middle jurassic unconformity’ consisting of the amalgamated sb 19–22. the log markers (a–m) recognised by michelsen (1989b) are defined by subtle changes in lithology, and are interpreted as chronostratigraphic markers that reflect basinwide changes in the relatively uniform and deep marine environment. pii hyllebjerg-1vedsted-1børglum-1frederikshavn-2 skagerrak–kattegat platform sorgenfrei–tornquist zone himmerland graben fj er ri ts le v fo rm at io n sb 15 k mfs 15 mfs 12 sb 12 ts 15 mfs 16 sb 16 sb 13 ts 13 m plu pii plu plu pii plu pii ji a/t børglum fault fj 5 fj 6 fj 7 fj 8 tst tst lst hst ne sw fj 4 fj 5 42 km 28 km 41 km k j l m i fii m em be r fiii m em be r fiv m em be r mfs 15 sb 15 ts 15 mfs 16 sb 16 ts 14 mfs 13 mfs 14 sb 14 lst hst tst tst hst lst tst hst gr sonic resspresspresspressp 3,4 100 m vedsted-1 frederikshavn-2 50 km sorgenfrei–tornquist zone børglum fault haldager fault fjerritslev fault danish basin skagerrak–kattegat platform hyllebjerg-1 børglum-1 ringkøbing–fyn high systems tracts lst lowstand tst transgressive hst highstand bounding surfaces sb sequence boundary mfs maximum marine flooding surface ts transgressive surface core fj 1 sequence depositional environments shoreface offshore stratigraphy lithostratigraphic boundary strata boundary (i–m) after michelsen (1989b) single dated samples top of chronostratigraphic unit, based on a series of ostracod analysis a/t aalenian or toarcian plu upper pliensbachian pii lower pliensbachian jl lower jurassic fig. 22. well-log panel showing the development of sequences fj 5 to fj 8 from hyllebjerg-1 to the more proximal well section (frederikshavn-2) on the skagerrak–kattegat platform. note the large amount of truncation at sb 15 which is supported biostratigraphically by ostracod faunas from the frederikshavn-2 core (fig. 15). flyvbjerg fm sb 20 ts 23 mfs 22 sb 23 sb 22 ts 22 sb 16 mfs 16 80 km 48 km 7 km 145 km 148 km 32 km 35 kmvemb-1 sp farsø-1 gr vedsted-1 sp terne-1 gr haldager-1 sp flyvbjerg-1 sp frederikshavn-2 sp skagen-2 sp børglum fm fjerritslev fm haldager fm ts 22 mfs 22 mfs 12 sb 23 ts 23 sb 22 sorgenfrei–tornquist zone danish basin skagerrak–kattegat platform sw ne fl 1 bø 1 ha 3 ha 2 ha 1 fj 10 fj 9 fj 8 fj 7 tst lst lst lst tst lst tst lst tst hst hst hst tst tst tst lst lst lst hst tst tst hst sb 19 sb 18 sb 16 sb 17 mfs 17 mfs 18 sb 21 ts 20 ts 19 mfs 19 mfs 20 sb 16 mfs 16 sb 21 sb 17 ts 21 mfs 17 1 2 fjerritslev fault børglum fault vemb-1 vedsted-1 skagen-2 terne-1 50 km sorgenfrei–tornquist zone børglum fault haldager fault fjerritslev fault danish basinringkøbing–fyn high skagerrak–kattegat platform flyvbjerg-1 haldager-1 100 m depositional environments fluvial estuarine lacustrine lagoonal shoreface offshore farsø-1 frederikshavn-2 systems tracts lst lowstand tst transgressive hst highstand bounding surfaces sb sequence boundary mfs maximum marine flooding surface ts transgressive surface core ha 1 sequence fig. 24. well-log panel showing the toarcian–kimmeridgian sequences across the danish basin to the skagerrak–kattegat platform. note the deep truncation at the base of the ‘middle jurassic’ (sb 19–22) outside the fault-bounded sorgenfrei–tornquist zone, clearly indicating that deposition continued in the sorgenfrei–tornquist zone in middle jurassic time while erosion prevailed elsewhere. the core coverage indicated on the frederikshavn-2 well represents a composite of cores from frederikshavn-1, -2 and -3. mfs 15 sb 19–22 mfs 22 sb 23 mfs 12 sb 15 sb 13 sb 12 mfs 11 sb 11 sb 9 mfs 7 l. cret. mfs 1 sb 5 børglum fault mfs 9 m fs 15sb 15 m fs 9 skagen-2 mejrup-1 sæby-1hyllebjerg-1 frederikshavn-2 ullerslev-1 børglum-1flyvbjerg-1 vemb-1 horsens-1 vedsted-1 ringkøbing–fyn high ts 19 mfs 22 mfs 16 sb 16 mfs 17 sb 17 mfs 16 vi 1 ha 1 ha 2 ha 3 fl 1 35 km 10 km34 km 15 km30 km65 km 41 km 92 km110 km 178 km sb 12 sb 9 sb 10 sb 11 sb 13 sb 14mfs 12 mfs 11 mfs 10 mfs 13 mfs 14 sb 16 sb 15 sb 14 sb 13 sb 12 sb 11 sb 10 sb 9 sb 5 sb 23 sb 22 sb 20 sb 21 sb 19 sb 17sb 5 sb 19–22 mfs 7 mfs 1 ts 1 fj 8 ga 1 fj 1 fj 2 fj 3 fj 4 fj 5 fj 6 fj 7 bø 1 fj 9 100 m depositional environments fluvial estuarine lacustrine lagoonal shoreface offshore 50 km sorgenfrei–tornquist zone børglum fault haldager fault fjerritslev fault danish basin skagerrak–kattegat platform skagen-2 frederikshavn-2 sæby-1 flyvbjerg-2 børglum-1 vedsted-1 hyllebjerg-1 mejrup-1 horsens-1 ullerslev-1 vemb-1 ringkøbing–fyn high bounding surfaces sb sequence boundary mfs maximum marine flooding surface ts transgressive surface ga 1 sequence fig. 25. well-log panel from the ringkøbing–fyn high (ullerslev-1) to the north-eastern basin margin (skagen-2) summarising the upper triassic – upper jurassic third-order sequences, their key surfaces and depositional environments. e2019430301-01 the amount of provenance information available for onshore and offshore sedimentary deposits in the north atlantic region is substantial and rapidly increasing. these data provide an improved understanding of reservoir geology (quality, diagenetic issues, regional source-to-sink relations and local stratigraphic correlations), and thereby can reduce hydrocarbon exploration risk. as such, the number of proprietary, industry-related and public research provenance studies has increased considerably in recent years, and the development and use of new analytical techniques has also caused a surge in the number of grains, isotopes and chemical elements analysed in each study. as a result, it is today close to impossible for the individual researcher or petroleum geologist to draw on all existing provenance data. and the vast expansion of data availability demands new and better methods to analyse and visualise large amounts of data in a systematic way. to this end, the geological survey of denmark and greenland (geus) and the norwegian petroleum directorate (npd) have established a web-based database of provenance data for the north atlantic area: the north atlantic provenance database. construction of the database was funded jointly by geus and npd. future maintenance and further development will be funded by the petroleum industry by subscription to the database. the database was launched in march 2019 and can be accessed at https://data.geus.dk/provenance. access to the database is granted via application to the steering committee, comprised of staff at geus, npd and sponsoring petroleum companies. the aims of the database are three-fold: • assemble provenance data from onshore and offshore greenland, faroe islands, norway and neighbouring areas. • make the data easily assessable to the petroleum industry and research institutions for visualisation and statistical analysis through a web application. • facilitate research as well as development of new provenance tools and techniques that can reduce exploration risk. the north atlantic provenance database: an introduction christian knudsen*1, martin sønderholm1, tjerk heijboer1, jeppe ågård kristensen2 and dag bering3 geus bulletin is free to individuals and institutions in electronic form. the author(s) retain copyright over the article contents. research article | open access geus bulletin vol 43 | e2019430301 | published online: 22 july 2019 https://doi.org/10.34194/geusb-201943-03-01 fig. 1. output map from the database showing the detrital zircon age distributions available as of july 2019. each data point (sample) is marked by either a yellow or a blue dot representing a sandstone sample or a present-day drainage sample, respectively. the rosette visualises the age distribution histogram where ‘north’ is 0 ma and ‘south’ is 1.8 ga with clockwise increasing ages. the length of each bar represents the frequency in 100 ma bins. inset: area shown in fig. 2. https://data.geus.dk/provenance https://doi.org/10.34194/geusb-201943-03-01 e2019430301-02 here, we provide a brief introduction to the database and its future development and expansion. we highlight the current capabilities with an example from east greenland. database contents, development and application the north atlantic provenance database is a spatiotemporal, object-rational database in postgresql. it consists of two main packages: • a database for storage of: – detrital zircon age data and other types of provenance data – metadata with information about the samples and the analytical procedure used • an analytical module allowing users to compile, compare and analyse the datasets. data at the time of publication (july 2019), the database consists of 1659 sandstone samples and 413 stream sediment samples, each containing detrital zircon u-pb ages, together with metadata such as coordinates, stratigraphic data, analytical data and a reference to where the data are published. the database contains more than 170 000 detrital zircon u-pb ages. the detrital zircon u-pb age data available in the database (july 2019) are summarised in fig.1. at present, detrital zircon u-pb age data constitute by far the main part of the existing provenance data in the north atlantic provenance database. other provenance data types, such as detrital rutile, monazite and apatite u-pb age data as well as stable heavy mineral distributions and heavy mineral chemistry, will be included at later stages. the analytical capabilities and possibilities are under constant change and the database structure is therefore flexible, to accommodate parameters that are not yet part of standard provenance analysis. most of the data from greenland that are stored in the database are comprised of onshore outcrop samples, but also include shallow core wells and samples from present-day drainage systems. in eastern greenland, c. 500 samples have been analysed for provenance properties. of these, more than 400 were analysed by geus. the database contains the age distribution of these detrital zircon grains as well as their heavy mineral distribution and compositions. in addition, the database also contains results from various studies n a f d e g c b 20 km fig. 2. geological map (1:500 000) of ymer ø in the kong oscar fjord area, east greenland (output directly from the database; location marked in fig.1). locations of stream sediment (blue) and sandstone samples (yellow) are marked, see fig. 1 for further details. inset: sampling locations of the seven samples (a–g; see fig. 3) discussed in the text. 0 1000 2000 3000 g f e d c b a ma 8307751 8105181 8105150 8105153 473730 473731 473729 n=58/135 n=121/154 n=115/140 n=109/143 n=101/148 n=79/121 n=37/54 fig. 3. detrital zircon age distribution histograms of the seven samples (a–g) selected in fig. 2. (rehnström et al. 2010). the histograms are generated according to thomsen et al. (2016). e2019430301-03 in east greenland: some are published (e.g. røhr et al. 2008; kirkland et al. 2009; sláma et al. 2011; olivarius et al. 2018), and others are to be found in unpublished theses and company reports. a wealth of offshore and onshore data is available for norway. some data have been extracted from published work, but the vast majority of the data are yet to be recovered from either unpublished academic work or company reports. on the norwegian shelf, a large number of samples have been analysed for their detrital zircon age distribution and heavy mineral content. again, some of these data are published (e.g. morton et al. 2005; lorenz et al. 2013; fleming et al. 2016) and others are yet to be extracted from unpublished theses and company reports. data from neighbouring territories such as west greenland, arctic russia and canada, the faroe islands, denmark and the uk are also accessible in the database. analytical module users can query the database via the analytical module interface, where they can view and analyse the relevant datasets for a given number of samples. from here, users can generate various visualisations of the data and export them as figures in pdf format. note that the data belong to the institutions that produced them and so they are not directly available to download from the database. however, the database contains full bibliographic information and links to where the data can be found. a more detailed description of the database architecture will be published alongside the database at a later date. here, we simply demonstrate some of functionality, using seven samples from gunner andersen land, ymer ø in kong oscar fjord, east greenland (rehnström et al. 2010; location in fig. 2). stream sediment samples from present-day drainage systems and proterozoic and devonian sandstones were chosen for further analysis. visual inspection of the detrital zircon age distribution rosettes (fig. 2) and histograms (fig. 3) show distinct modal variation between the seven samples. this is confirmed by the kolmogorov-smirnov (ks) dissimilarity test (fig. 4 and table 1). some of the metadata available in the database for these samples are shown in table 2. sample 473729 (fig. 3a) represents the sandertop formation of the upper proterozoic lyell land group, part of the eleonore bay supergroup. detrital zircon ages span the palaeoproterozoic–mesoproterozoic eras, with a few archaean grains. samples 473730 (fig 3c) and 473731 (fig. 3b) represent the devonian kap kolthoff group consisting of immature sandstone and conglomerate. the detrital zircon age distributions are dominated by a palaeoproterozoic a b fig. 4. kolmogorov-smirnov (ks) dissimilarities of the seven selected samples. a: cumulative age distributions. b: hierarchical clustering. table 1. kolmogorov-smirnov (ks) dissimilarity matrix of the seven selected samples. x473729 x473730 x473731 x8105150 x8105153 x8105181 x8307751 x473729 0.631 0.585 0.190 0.126 0.344 0.268 x473730 0.631 0.064 0.647 0.631 0.328 0.620 x473731 0.585 0.064 0.622 0.610 0.330 0.589 x8105150 0.190 0.647 0.622 0.176 0.369 0.127 x8105153 0.126 0.631 0.610 0.176 0.343 0.163 x8105181 0.344 0.328 0.330 0.369 0.343 0.330 x8307751 0.268 0.620 0.589 0.127 0.163 0.330 e2019430301-04 peak around 1980 ma, which is well known from the tonalitic gneiss that occur in the basement northeast of the area (kalsbeek et al. 2008). samples 8105153, 8105150 and 8307751 (figs 3d, e, f) represent stream sediments collected from the present-day drainage system on ymer ø. they have very similar detrital zircon age distributions that resemble those of the upper proterozoic eleonore bay supergroup bedrock (see sample 473729; fig 3a). this supports the view that stream sediment offers a good representation of the catchment bedrock geology. stream sample 8105181 (fig. 3g) from the mainland has a very complex detrital zircon age distribution suggesting a fundamental difference in the bedrock geology in the catchment area of this sample. the analytical module allows users to calculate and visualise cumulative age distributions and kolmogorov-smirnov dissimilarities (fig. 4) – a widely used method for comparing mineral age distributions. this is enabled through a plugin between the database and the freely available statistical programming software r (r development core team 2008). these functionalities are part of the ‘provenance’ package’ developed specifically for detrital sediment provenance analysis (vermeesch et al. 2016). the similarities between e.g. samples 473730 and 473731 are clearly seen in figures 4a and 4b, and in table 1. outlook compiling the large amount of available provenance data into a regional, cross-border, web-database will make these types of data much more accessible and applicable to industry and the research community. in doing so, we hope to promote the use of these data in studies of the north atlantic region. with an extensive database covering both sides of the atlantic ocean more comprehensive source-to-sink analyses can be made, resulting in an improved understanding of onshore–offshore provenance relationships. in the long-term, we hope to include more data from neighbouring geographic areas such as the russian and canadian arctic, since detrital material in the north atlantic may have been derived from these areas. it is envisaged that more than 100 000 detrital zircon grains from the north atlantic region have been dated, and the aim is to capture the majority of these in the database, making them available for data comparison (e.g. with statistical tools) and visualisation to enhance the understanding of the regional reservoir geology. users are expected to upload their own provenance data in return for using the database. it is also possible to restrict access to certain data in the database, so that they are kept confidential for a time. this is an important feature for these types of datasets, many of which are funded by private companies and have confidentiality clauses imposed for a finite period. for more information contact the lead-author. references fleming, e.j., flowerdew, m.j., smyth, h.r., scott, r.a., morton, a.c., omma, j.e., frei, d. & whitehouse m. j. 2016: provenance of triassic sandstones on the southwest barents shelf and the implication for sediment dispersal patterns in the northwest pangea. marine and petroleum geology 78, 516–535. https://doi.org/10.1016/j.marpetgeo.2016.10.005 kalsbeek, f., thrane, k., higgins, a.k., jepsen, h.f., leslie, a.g., nutman, a.p. & frei, r. 2008: polyorogenic history of the east greenland caledonides. in: higgins et al. (eds): memoir 202: the greenland caledonides: evolution of the northeast margin of laurentia. boulder, colorado: geological society of america, 55–72. https://doi. org/10.1130/2008.1202(03) kirkland, c.l., pease, v., whitehouse, m.j. & ineson, j.r. 2009: provenance record from mesoproterozoic-cambrian sediments of peary land, north greenland: implications for the ice-covered greenland shield and laurentian paleogeography. precambrian research 170, 43–60. https://doi.org/10.1016/j.precamres.2008.11.006 lorenz, h., gee, d.g., korago, e., kovaleva, g., mcclelland, w.c., gilotti, j.a. & frei, d. 2013: detrital zircon geochronology of palaeozoic novaya zemlya – a key to understanding the basement of the barents shelf. terra nova 25, 496–503. https://doi.org/10.1111/ter.12064 morton a.c., whitham, a.g. & fanning c.m. 2005: provenance of late cretaceous to palaeocene submarine fan sandstones in the norwegian sea: integration of heavy mineral, mineral chemical and zircon age data. sedimentary geology 182, 3–28. https://doi.org/10.1016/j. sedgeo.2005.08.007 olivarius, m., bjerager, m., keulen, n., knudsen, c. & kokfelt, t.f. 2018: provenance of basinal sandstones in the upper jurassic hareelv formation, jameson land basin, east greenland. in: ineson, j. & bojesentable 2. example metadata available for the seven samples in fig. 2 storage sample type lithology lithostratigraphy base age top age reference number 473729 rock sample sandstone lyell land group sandertop formation tonian tonian rehnstrøm et al. 2010 473730 rock sample sandstone kap kolthoff group givetian famennian rehnstrøm et al. 2010 473731 rock sample sandstone kap kolthoff group givetian famennian rehnstrøm et al. 2010 8105150 stream sediment sample rehnstrøm et al. 2010 8105153 stream sediment sample rehnstrøm et al. 2010 8105181 stream sediment sample rehnstrøm et al. 2010 8307751 stream sediment sample rehnstrøm et al. 2010 https://doi.org/10.1016/j.marpetgeo.2016.10.005 https://doi.org/10.1016/j.marpetgeo.2016.10.005 https://doi.org/10.1130/2008.1202(03) https://doi.org/10.1130/2008.1202(03) https://doi.org/10.1016/j.precamres.2008.11.006 https://doi.org/10.1111/ter.12064 https://doi.org/10.1016/j.sedgeo.2005.08.007 https://doi.org/10.1016/j.sedgeo.2005.08.007 e2019430301-05 koefoed, j.a. (eds): petroleum geology of the upper jurassic – lower cretaceous of east and north-east greenland: blokelv-1 borehole, jameson land basin. geological survey of denmark and greenland bulletin 42, 39–64. r development core team 2008. r: a language and environment for statistical computing. r foundation for statistical computing, vienna, austria. isbn 3-900051-07-0, url http://www.r-project.org. rehnström, e.f., thrane, k., kokfelt, t.f. & frei, d. 2010: age distribution of detrital zircon grains in sandstones and stream sediments from east greenland north of 70°n. geological survey of denmark and greenland report 2010/130. 125 pp røhr, t. s., andersen, t. & dypvik, h. 2008: provenance of lower cretaceous sediments in the wandel sea basin, north greenland. journal of the geological society 165, 755–767. https://doi.org/10.1144/001676492007-102 sláma, j., walderhaug, o., fonneland, h., kosler, j. & pedersen, r.b. 2011: provenance of neoproterozoic to upper cretaceous sedimentary rocks, eastern greenland: implications for recognizing the sources of sediments in the norwegian sea. sedimentary geology 238, 254–267. https://doi.org/10.1016/j.sedgeo.2011.04.018 thomsen, t.b., heijboer, t., & guarnieri, p. 2016: jagedisplay: software for evaluation of data distributions in u-th-pb geochronology. geological survey of denmark and greenland bulletin 35, 103–106. vermeesch, p., resentini, a. & garzanti, e. 2016: an r package for statistical provenance analysis. sedimentary geology 336, 14–25. https:// doi.org/10.1016/j.sedgeo.2016.01.009 how to cite knudsen, c., sønderholm, m., heijboer, t., kristensen, j.å. & bering, d. 2019: the north atlantic provenance database: an introduction. geological survey of denmark and greenland bulletin 43, e2019430301. https://doi.org/10.34194/geusb-201943-03-01 *corresponding author: christian knudsen | e-mail: ckn@geus.dk 1 geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350, copenhagen k, denmark. 2 department of physical geography and ecosystem science, lund university, sölvegatan 12, s-223 62 lund, sweden. 3 norwegian petroleum directorate, professor olav hansens vei 10, postboks 600, 4003 stavanger, norway. http://www.r-project.org https://doi.org/10.1144/0016-76492007-102 https://doi.org/10.1144/0016-76492007-102 https://doi.org/10.1016/j.sedgeo.2011.04.018 https://doi.org/10.1016/j.sedgeo.2016.01.009 https://doi.org/10.1016/j.sedgeo.2016.01.009 https://doi.org/10.34194/geusb-201943-03-01 mailto:ckn%40geus.dk?subject= geological survey of denmark and greenland bulletin 6, 19-29 19 the fluviatile bristol elv formation, a new middle jurassic lithostratigraphicunit fromtraillø,north-east greenland jens therkelsen and finn surlyk a new lithostratigraphic unit, the bristol elv formation, is erected in this paper. it is only known from traill ø, east greenland, where it unconformably overlies triassic redbeds of the fleming fjord formation and is overlain by lithologically similar shallow marine upper bajocian sandstones of the pelion formation. the age of the formation is not well constrained but is probably early bajocian. the bristol elv formation is at least 155 m thick and consists of conglomerates, coarse-grained pebbly sandstones and subordinate mudstones, deposited in braided rivers. a finer-grained lacustrine/floodplain unit, c. 37 m thick, is interbedded with the fluvial sandstones at one locality. deposition of the fluvio-lacustrine bristol elv formation marks a major change in basin configuration and drainage patterns, reflecting the onset of the important, protracted middle–late jurassic rift event in east greenland. keywords: fluvial, lacustrine, middle jurassic sediments, new formation, north-east greenland j.t., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. present address: skude & jacobsen, næstvedvej 1, dk-4760 vordingborg, denmark. e-mail: jth@sjas.dk f.s., geological institute, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. geological survey of denmark and greenland bulletin 5, 19–29 © geus, 2004 the first detailed description of the jurassic sandstones on the islands of traill ø and geographical society ø (fig. 1) was presented by donovan (1953, 1955, 1957) who grouped the deposits in the yellow series of maync (1947). donovan’s work was focused on the sediments exposed in the bjørnedal area in south-eastern traill ø while the sandstones exposed at mols bjerge and svinhufvud bjerge received less attention. donovan (1953, p. 64) suggested that black shales interbedded with sandstones and occasional conglomerates, which he termed the plant beds, were deposited in a marine embayment, or possibly on the subaerial part of a debris fan. higher in the sandstone unit, he found evidence for periodic marine incursions as indicated by occasional ammonite-bearing levels. a lithostratigraphic scheme covering the jurassic of jameson land in east greenland was erected by surlyk et al. (1973) and was extended to the areas north of kong oscar fjord by surlyk (1977). the middle jurassic sandstones of traill ø and geographical society ø were placed in the pelion member of the vardekløft formation. this scheme has recently been revised in the light of much new work in the region resulting in rank changes and establishment of new formations and members (see surlyk 2003, fig. 5). fieldwork in the traill ø area has revealed that sandstones formerly referred exclusively to the shallow marine pelion formation (pelion member in: surlyk 1977) actually consist of a lower fluvial unit overlain by marine sandstones (price & whitham 1997; stemmerik et al. 1997). the fluviatile deposits are placed in a new lithostratigraphic unit, the bristol elv formation, which is erected here as the basal unit of the middle jurassic vardekløft group on traill ø (fig. 2). the new formation consists dominantly of conglomerates, coarse-grained pebbly sandstones and subordinate mudstones and was deposited in a braided river environment, probably in middle jurassic, early bajogeus bulletin no 5.pmd 29-10-2004, 11:1419 20 72 ■ ■ ■ ■ ■ ■ geographical society ø traill ø 24ºw 23ºw 22ºw 73º00'n 72º45'n 72º30'n 72º15'n 72 00′ n svinhufvud bjerge 10 km mols bjerge vælddal jameson land ■ ■ ■ ■ ■ ■ månedal ■ ★ ★ ★ ★ kong oscar fjord n=16 n=22 n=31 n=23 2 1 4 3 ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ g re en la nd localities faults palaeocurrent direction (n=22: number of measurements) distribution of the bristol elv formation ★1 ■ cian time. in one section, fine-grained sandstones, mudstones, mediumto coarse-grained sandstone beds and thin coal seams occur intercalated with the fluvial sandstones, and are interpreted as lacustrine or floodplain deposits. bristol elv formation new formation history. the strata assigned here to the new bristol elv formation were included in the yellow series (maync 1947) by donovan (1953) and were referred to the pelion member of the vardekløft formation by surlyk (1977). the fluvial nature of the lower part of the succession was recognised independently by price & whitham (1997) and stemmerik et al. (1997), and labelled pm1 by the former authors. fig. 1. map showing the distribution of the bristol elv formation in the traill ø area, the position of localities and palaeocurrent directions. the map is modified from stemmerik et al. (1997). geus bulletin no 5.pmd 29-10-2004, 11:1420 21 name. after the river bristol elv in the southern part of mols bjerge, traill ø (fig. 1). type section. southern part of svinhufvud bjerge (fig. 1, locality 1) on the south coast of traill ø (figs 1, 3). reference sections. north-eastern svinhufvud bjerge (fig. 1, locality 2), northern and southern mols bjerge (locations 3, 4) and vælddal, all on traill ø (figs 1, 4). thickness. the formation is at least 155 m thick in the southern part of svinhufvud bjerge, whereas thicknesses in excess of 80 m are recorded in northern mols bjerge and at bristol elv in the southern part of mols bjerge. thicknesses of 280 m and 520 m in the southern and northern svinhufvud bjerge areas, respectively, and 310 m in northern mols bjerge were stated by price & whitham (1997), but these large values have not been corroborated by our study (surlyk & noenygaard 2001). lithology. the bristol elv formation consists mainly of conglomerates, pebbly sandstones and sandstones in the mols bjerge area, whereas more fine-grained deposits are intercalated in the svinhufvud bjerge area. the bulk of the formation consists of yellowish to whitish, poorly sorted, coarse-grained sandstones, pebbly sandstones and fine pebble conglomerates with subrounded to well-rounded quartzite pebbles and mudstone intraclasts. black to dark grey mudstones with centimetre-thick autochthonous coal beds occur intercalated with the sandstones and conglomerates at several levels. the coarse-grained deposits show largescale trough cross-bedding, but the structures are mostly poorly defined. the cross-bedded sets are 0.15–2.0 m thick, and form cosets up to 7 m thick. also observed are large-scale scour-and-fill structures (fig. 5), planar cross-bedding, planar lamination, rare water escape structures and imprints of tree trunks, which may be found in accumulations. the sediments commonly form fining-upwards units, up to 16 m thick. on the south coast of traill ø, a sandstone unit, c. 19 m thick, shows evidence of gently dipping bedforms (fig. 6, corresponding to the 134–153 m interval in fig. 3). at svinhufvud bjerge, the coarse-grained units are separated by mudstone units up to 6 m thick, while the coarsegrained units at mols bjerge are amalgamated without any fine-grained interbeds. the mudstones in the svinhufvud bjerge area are very uniform in grain size and overlie the sandstones with a sharp boundary. they show a well-developed to faint lamination, which is commonly disturbed by rooting. a few centimetre-thick layers of mediumgrained sandstone occur in the upper part of the mudstone units at the same level as the rootlets. the orientation of cross-bed trough axes and foreset azimuths of planar cross-beds in the sandstones at mols bjerge generally indicates westerly palaeocurrent directions, whereas the directions in the svinhufvud bjerge area are generally towards the south-east. stylolites are common in the coarse-grained sandstones and fine-grained conglomerates at mols bjerge but have not been observed at svinhufvud bjerge. a succession of mainly black to dark grey mudstones and fine-grained sandstones, c. 37 m thick, with intercalated coal beds and coarse-grained sandstones occurs in the type section at svinhufvud bjerge (fig. 3, 170–207 m in log). the sandstones are trough crossbedded, structureless, cross-laminated or lenticular fleming fjordscoresby land bristol elv pelion hiatus fossilbjerget vardekløft parnas mb olympen bernbjerg formation hall bredning kimmeridgian oxfordian callovian bathonianm id dl e u pp er ju ra ss ic tr ia ss ic bajocian groupstagesystem u u m l m l u u m l l fig. 2. lithostratigraphic scheme of the jurassic succession in the traill ø area. the age of the bristol elv formation is poorly constrained, but an early bajocian age is the most likely. based on clemmensen (1980) and d. strogen (personal communication 2000); see also surlyk (2003). geus bulletin no 5.pmd 29-10-2004, 11:1421 22 30 20 10 0 320 310 300 290 280 270 260 250 240 220 210 200 190 180 170 160 150 140 130 120 110 mud sand pebble c. 19 m poorly exposed c. 40 m poorly exposed c. 30 m basaltic sill m br is to l el v f or m at io n ? tr ia ss ic pe lio n f or m at io n fl em in g fj or d fo rm at io n mud sand pebble m mudstone sandstone pebbly sandstone conglomerate coal lithology palaeocurrent direction degree of bioturbation belemnite rootlets stems/logs plant fragments stylolites coalified logs diplocraterion habichi ophiomorpha nodosa coarsening-upwards fining-upwards sub-vertical burrows structureless sand structureless conglomerate parallel lamination wavy bedding cross-lamination structures tabular cross-bedding trough cross-bedding lenticular bedding concretion mudstone clasts slump hummocky and swaley cross-stratification geus bulletin no 5.pmd 29-10-2004, 11:1422 23 facing page: fig. 3. type section (locality 1) of the bristol elv formation. south svinhufvud bjerge, traill ø. position shown in fig. 1. bedded, and display both sheet-like and lenticular geometries. some of the sandstones show poorly developed wavy surfaces, are weakly bioturbated and contain rare mudstone flasers. the mudstones are mainly laminated or weakly laminated and contain conspicuous root horizons, in situ tree stumps, fern leaves and early diagenetic sideritic concretions. sandstone beds up to 10 cm thick with wave ripple crossstratification, similar to micro-hummocky cross-stratification, occur in the mudstones. the ripples have wavelengths up to 30 cm and heights up to 5 cm. immediately above one of these sandstone beds, rounded pebbles up to 4 cm in diameter are present. a sandstone bed, c. 1 m thick, showing swaley cross-stratification occurs intercalated in the mudstones (fig. 3, 172–173 m in log, fig. 7b). the wave-rippled and swaley cross-stratified beds form part of two small coarsening-upwards units, 7 m and 5 m thick, in the lowest part of the succession (fig. 3, 170–182 m in log, fig. 7a). the lower part of the units consists of laminated and faintly laminated mudstones, which in the lowest unit are intercalated with wave-rippled and swaley cross-stratified sandstone beds. the units grade upwards into very fine-grained sandstones, showing poorly developed wavy surfaces, cross-lamination, lenticular bedding and rare mudstone flasers. unidentified trace fossils occur in both units and the top beds in the upper coarsening-upwards unit are penetrated by rootlets. above these two units, coal beds up to 0.45 m thick and thin mudstone beds occur together with 0.2–2.5 m thick beds of trough cross-bedded, medium-grained sandstone to fine pebble conglomerate with sharp basal surfaces, showing palaeocurrents towards the east-south-east (fig. 3, 182– 207 m in log). locally, these beds are overlain by fineto very fine-grained sandstones, forming poorly defined fining-upwards successions. pyrite is not present in the coals, but occasionally replaces organic detritus in the sandstones. the coals consist almost entirely of non-detrital vitrinite, which together with the presence of root horizons beneath the coal beds show that they are autochthonous. boundaries. the formation unconformably overlies the upper triassic fleming fjord formation at svinhufvud bjerge and mols bjerge as well as in vælddal (clemmensen 1980; d. strogen, personal communication 2000). the upper boundary is placed at an erosional surface draped by a pebble lag overlain by marine sandstones of the pelion formation, which is dominated by mediumto coarse-grained, planar cross-bedded and structureless sandstones with ammonites, belemnites, bivalves and marine trace fossils. distribution. the formation is known only from svinhufvud bjerge, mols bjerge and vælddal in eastern traill ø (fig. 1). the lacustrine/backswamp unit described from the type section is probably correlative to the plant beds of donovan (1953, 1957), which occur in vælddal (henrik vosgerau, personal communication 1998). geological age. no macrofaunas were found within the bristol elv formation. a few relatively well preserved but as yet unidentified fern leaves were retrieved from a single bed in the lacustrine/floodplain unit in the southern part of svinhufvud bjerge. harris (1946) reported a stem identified as equisetites sp. a. of inferred early jurassic age from kap palander in the northernmost mols bjerge. harris (1946) pointed out, however, that this specimen also resembles species of late triassic or middle jurassic ages, and that it does not give any precise age indication. preliminary palynological analysis of samples from the lacustrine/backswamp deposits suggests a broad late toarcian – bathonian age (karen dybkjær, personal communication 1998). age diagnostic palynomorphs include callialasporites dampieri (late early toarcian or younger), callialasporites turbatus (late toarcian or younger), callialasporites segmentatus (late early toarcian or younger) and foraminisporis jurassicus (middle rhaetian – bajocian) (batten & koppelhus 1996). the upper bajocian cranocephalites borealis zone is represented in the immediately overlying pelion formation sandstones (donovan 1953, 1957; callomon 1993; alsen 1998). this is the lowest middle jurassic ammonite zone recognised in east greenland. in jameson land, the pelion formation contains a relatively thick marine sandstone unit without ammonites below the lowest occurrence of cranocephalites borealis. this unit overlies dark mudstones of the sortehat formation the top of which is of early bajocian age (underhill & partington 1993; koppelhus & hansen 2003). the unit is thought to be a distal marine correlative of the bristol elv formation. this is supported by the stratigraphic position of both units, underlying marine pelion sandstones of the c. borealis chronozone geus bulletin no 5.pmd 29-10-2004, 11:1423 24 0 10 20 30 40 50 60 70 m pebbl.sand pebbl.sandpebbl.sand 0 10 20 30 40 50 60 70 m locality 2 north-eastern svinhufvud bjerge locality 3 northern mols bjerge locality 4 southern mols bjerge 30ð40 m 0 5 10 15 20 25 m geus bulletin no 5.pmd 29-10-2004, 11:1424 25 and by the marked lithological similarity of the pelion and bristol elv formations. lower jurassic rocks have never been documented outside jameson land and all available data thus point towards an early bajocian age for the bristol elv formation. depositional environment the coarse-grain size, the presence of fining-upwards units and trough cross-bedding, the unidirectional palaeocurrents towards the west (in mols bjerge) and south-east (in svinhufvud bjerge) and the abundance of mudstone intraclasts, the absence of dinoflagellate cysts and marine body and trace fossils indicate that the main part of the bristol elv formation was deposited in a high-energy fluvial environment. studied samples all show very low total sulphur (ts) contents (max. 0.43%) and high c/s values (> 10). these data support the interpretation of a terrestrial environment of deposition (berner & raiswell 1984). in the svinhufvud bjerge area, one sandstone unit, approximately 20 m thick, shows macroforms interpreted as downstream or lateral accretion structures such as epsilon cross-bedding (allen 1963; fig. 6). the lower c. 13 m of this unit displays a fining-upwards trend, possibly representing lateral channel migration (allen 1964, 1965). the depth of the channel corresponds to at least the thickness of the macroform (leeder 1973). if the interpretation of the macroforms is correct, the channel depths were in the 10–13 m range. fluvial styles and models were categorised in terms of channel sinuosity/braiding, sediment type and characteristic architectural elements by miall (1985, 1996). the depositional features of the bristol elv formation, especially concerning the within-channel element and the relatively large channel depths suggest deposition in the ‘perennial deep braided river’ type of miall (1996). reliable interpretation of a fluvial system cannot, however, be based on vertical sections alone (miall 1996). analysis of bounding surfaces, and extent, shape and facies relations of the architectural elements has, however, not been possible in the present case due to the nature of the outcrop. the interbedded mudstones may, due to the sharp boundary to the underlying coarse sandstones and conglomerates, represent relatively abrupt channel abandonment and subsequent passive in-filling of channels and thus cover a confined area, or they may be more extensive bodies covering the entire floodplain area. the studied exposures do not allow conclusions on the lateral extent of the mudstone units. the finegrained homogeneous nature of the mudstones, however, suggests that the distance to the nearest active fluvial channel must have been relatively large leading to deposition of only the finest grain sizes. the sudden change from very coarse sandstone to homogeneous mudstone also indicates a very abrupt abandonn s fig. 5. coarse-grained sandstone and fine-grained conglomerate with large-scale scour-and-fill structures (indicated by dashed lines below encircled person). bristol elv formation; north-eastern svinhufvud bjerge. facing page: fig. 4. composite detailed reference section from northeastern svinhufvud bjerge (locality 2) and reference sections from northern and southern mols bjerge (localities 3, 4). positions shown in fig. 1. for legend, see fig. 3. geus bulletin no 5.pmd 29-10-2004, 11:1425 26 ment of the active channels, which in a relatively short time shifted to a position farther away. floodplains in braided river systems are not commonly described in the literature, but studies by reinfelds & nanson (1993) of the waimakariri river, new zealand, show that braided rivers, contrary to the common conception, may include large areas of finegrained floodplains. during avulsion events, extensive wetland areas were established and contributed greatly to trapping and deposition of large volumes of finegrained material (e.g. smith et al. 1989). a similar situation may have occurred at least twice during the lifetime of the bristol elv formation river system, as shown by the occurrence of 3–6 m thick floodplain mudrocks in the section on the south coast of traill ø (figs 1, 6). the presence of root horizons and thin coal beds shows that the floodplain was densely vegetated and developed into peat swamps, before the river channel migrated back over the area and peat formation ceased. fine-grained sediment present in the middle part of the type section in svinhufvud bjerge is interpreted to have been deposited in a lacustrine/backswamp environment. fine-grained sandstones intercalated with black and dark grey mudstone in the lower part of this lacustrine/backswamp unit show swaley cross-stratification and wave ripple cross-lamination associated with a pebble lag (figs 3, 7). swaley cross-stratification is interpreted as the result of storm-wave deposition above fairweather wave base and has been described from very shallow depths in the large lake superior (greenwood & sherman 1986; sherman & greenwood 1989). the wave ripple cross-lamination shows that some vigorous agitation must have occurred, probably during storm events (e.g. allen 1982). the structures and the associated pebble lag are probably the result of shoreface erosion during storms, and subsequent transport into deeper waters by storm-induced currents (cf. dam & surlyk 1992, 1993). in the uppermost part of the lacustrine succession, the occurrence of a unit that shows coarsening-upwards from mudstone to very finew e 20 m fig. 6. fluvial, floodplain and lacustrine deposits on the south coast of traill ø (locality 1). note the 20 m thick sandstone body in the centre (see fig. 3, 134–153 m in log) showing down-stream or lateral accretion structures (dashed lines) with bedforms dipping gently to the right (east). the sandstone body overlies dark floodplain mudstones. type section of the bristol elv formation. geus bulletin no 5.pmd 29-10-2004, 11:1426 27 grained sandstone (fig. 3, 177–182 m) possibly records in-filling of the lake, resulting in lake shoreline progradation and gradual shallowing. the presence of rootlets shows that the lake was sufficiently shallow to allow colonisation of vegetation. higher in the succession (fig. 3, 182–207 m), abundant conspicuous root horizons show that vegetation spread across the shores of the lake, which eventually turned into a backswamp environment. the repeated development of autochthonous coal beds in the middle and upper part of the unit suggests that the backswamps were densely vegetated. the 0.2–2.5 m thick trough cross-bedded sandstone beds associated with the coal beds, probably represent splays into the backswamp from active channels, which were possibly situated north-west of the area. the lack of pyrite in the coal is a good indication of deposition in a freshwater environment (cohen et al. 1984; brown & cohen 1995; phillips & bustin 1996). the plant beds of donovan (1953, 1957) in vælddal are interbedded with coarse-grained sandstones and conglomerates dominated by large-scale trough crossbedding, while hummocky and/or swaley cross-stratification, small-scale cross-lamination, root horizons and fossilised leaves occur in the finer grained sediments (henrik vosgerau, personal communication 1998). if the plant beds of donovan (1953, 1957) are correlatives of the mudstone-dominated part of the bristol elv formation typesection in svinhufvud bjerge, a system with a b fig. 7. lacustrine deposits from the bristol elv formation, south coast of traill ø (locality 1). length of knife is 21 cm. a: thin micro-hummocky cross-stratified / wave ripple cross-laminated sandstone intercalated with laminated mudstone. b: sandstone bed showing swaley crossstratification. geus bulletin no 5.pmd 29-10-2004, 11:1427 28 scattered lakes may have existed in the area. the lateral extent of this system must have been at least 20 km. gradual in-filling of the lake resulted in development of a wetland area with peat swamps, which was periodically covered by sheet-sands and cut by confined channels where sand was deposited. thin coarsening-upwards units, which probably represent crevasse splays or deltas, show that active fluvial channels were present in the adjacent area, and the lake system was eventually replaced by a fluvial braided channel system, represented by trough cross-bedded coarse-grained pebbly sandstones. this environment persisted until the area was transgressed by the sea and the shallow marine sandstones of the pelion formation were deposited. acknowledgements this study was undertaken under the auspices of the project ‘resources of the sedimentary basins of north and east greenland’, supported by the danish research council. we thank karen dybkjær and stefan piasecki for palynological contributions, jan andsbjerg, gregers dam, and jon r. ineson for critical reading of an early manuscript version and reviewers d. strogen and michael larsen for constructive reviews. references allen, j.r.l. 1963: the classification of cross-stratified units, with notes on their origin. sedimentology 2, 93–114. allen, j.r.l. 1964: studies in fluviatile sedimentation: six cyclothems from the lower old red sandstone, anglo-welsh basin. sedimentology 3, 163–198. allen, j.r.l. 1965: a review of the origin and characteristics of recent alluvial sediments. sedimentology 5, 89–191. allen, j.r.l. 1982: sedimentary structures; 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(eds): siliciclastic sequence stratigraphy: recent developments and applications. american association of petroleum geologists memoir 58, 449–484. geus bulletin no 5.pmd 29-10-2004, 11:1429 geological survey of denmark and greenland bulletin 28, 2013, 49-52 49 geochemistry and petrology of gold-bearing hydrothermal alteration zones on qilanngaarsuit, southern west greenland martin koppelberg, annika dziggel, denis martin schlatter, jochen kolb and franz michael meyer during field work in 2008, the geological survey of denmark and greenland investigated the gold potential of hydrothermal vein systems in the nuuk region of the archaean north atlantic craton. a new gold occurrence was discovered on the island of qilanngaarsuit, 35 km south of nuuk. two cross sections through hydrothermal alteration zones, that locally contain up to 672 ppb au, were mapped and sampled in detail. in this study, we present petrological and geochemical data in order to constrain the conditions for ore formation and transport of elements during fluid–rock interaction. geological setting qilanngaarsuit in southern west greenland (fig. 1) is situated in the godthåbsfjord gold province, a c. 20 km by 150 km wide, ne-trending sector along the ivinnguit fault system. several gold occurrences such as storø, qussuk, bjørneøen and sw isua have been described from this sector (garde et al. 2012; scherstén et al. 2012; kolb et al. 2013). the ivinnguit fault, situated north of the island, is a deep-crustal shear zone, which formed during terrane amalgamation and it represents the border between the færingehavn and akia terranes (nutman & friend 2007). qilanngaarsuit island is dominated by eoarchaean tonalite-trondhjemite-granodiorite (ttg) gneisses of the færingehavn terrane, which are overlain by amphibolites and aluminous cordierite-orthoamphibole gneisses that originated from c. 2840 ma old protoliths (nutman & friend 2007). four deformation events can be distinguished, involving north-vergent thrusting, isoclinal folding and the formation of late, upright, open to tight folds (e.g. kolb et al. 2013). the gold-bearing quartz veins are hosted by amphibolites in the central part of the island and surrounded by approx. 8 m wide hydrothermal alteration zones. the foliation-parallel, c. 10–20 cm wide, quartz veins can be followed over several hundred metres along strike. structural data indicate that they formed in response to flexural, slip folding during the late-tectonic evolution of the region (kolb et al. 2013). one sample from a hydrothermal alteration zone surrounding the veins contains up to 672 ppb au and several other vein and alteration zone samples have elevated au contents (> 20 ppb au; kolb et al. 2009). © 2013 geus. geological survey of denmark and greenland bulletin 28, 49–52. open access: www.geus.dk/publications/bull 50°w 64°n 63°n s q b 50 km eoarchaean gneiss anorthosite-gabbro complexes orthogneiss and granitic rocks (mesoto neoarchaean) tonalitic to granodioritic plutons (mesoto neoarchaean) granite (mesoto neoarchaean) quaternary cover ice cover granulite facies qôrqut granite complex fault qilanngaarsuit fisk efjord fau lt go dth åb sfj ord fiske fjord ameralik ameralik serm ilik bjørnesund block færingehavn terrane tasiusarsuaq terrane akia terrane tre brødre terrane tre brødre terrane isukasia terrane bu kse fjord bjørn esu nd grædefjord godthåbsfjord– ameralik belt ivinnguit fault sermilik block nuuk færingehavn kapisilik terrane terrane or block boundary supracrustal belts (undifferentiated) greenland fig. 1. geological map of the nuuk region (modified from allaart 1984). b: bjørneøen, q: qussuk, s: storø. 5050 petrology and geochemistry two profiles from the unaltered wall rocks through the hydrothermal alteration zones were investigated (fig. 2). two types of amphibolite can be distinguished: homogeneous amphibolite in the footwall and compositionally layered amphibolite in the hanging wall of the hydrothermal vein system. their protoliths were low-k tholeiites depleted in light rare-earth elements (lree), geochemically similar to other amphibolites in the buksefjord region (chadwick 1981). generally, the amphibolites are fineto medium-grained and consist of hornblende (40 vol.%), plagioclase (30  vol.%), clinopyroxene (20   vol.%) and clinozoisite/zoisite/epidote (10  vol.%). metamorphic garnet is locally present in the layered amphibolite. retrogression is indicated by the transformation of plagioclase to a fine-grained assemblage of zoisite and quartz, and by the replacement of amphibolite and clinopyroxene by epidote and clinozoisite. within the alteration zone, the amphibolite-facies mineral assemblages are replaced by a high temperature alteration assemblage of garnet, quartz, plagioclase, biotite, and sillimanite (figs 2, 3). in contact with the veins, the hydrothermal alteration zone consists of up to 50  vol.% garnet, 15  vol.% plagioclase, 15  vol.% quartz, 10  vol.% biotite and 10   vol.% sillimanite (fig. 2). relict amphibole facies minerals such as hornblende and clinopyroxene are locally preserved, indicating that these minerals formed during regional metamorphism prior to the mineralisation. ore minerals make up ≤1 vol. % of the rocks; they include pyrite, pyrrhotite and chalcopyrite. mass-balance calculations based on whole-rock, majorand traceelement data (kolb et al. 2009; koppelberg 2011) and using the method of gresens (1967) indicate that the ore fluid was enriched in si, k, lree, au, cr, cu, zn, mo and as (fig. 4). in si-rich, vein-dominated samples, the hydrothermal overprint was associated with a volume increase of 14–62%. in contrast, the sillimanitebearing samples record a significant volume loss of 15–50%, and are depleted in si (fig. 4). this suggests that at least some of the si in the quartz veins was leached from the surrounding wall rocks. garnet in the alteration zones is rich in almandine (koppelberg 2011). most grains are essentially unzoned and have fig. 2. lithological logs of profiles a and b (modified from schlatter 2009). amphibolite with weak grt-bt alteration (pl (xab0.22-0.45), hbl, qtz, cpx, bt, grt (alm46-62, prp20-39,grs80-12) unaltered amphibolite (hbl, pl, cpx, qtz, ttn) retrogressed amphibolite (pl (xab0.49-0.61), hbl, zo, cpx ± qtz, ttn, ep, bt, cal unaltered amphibolite with grt (alm53-59,prp13-20, grs20-22), hbl, pl (xab0.19-0.29), cpx, qtz grt-bt alteration, vein-dominated, without sil: pl (xab0.59-0.65), qtz, grt (alm53-64-prp10-33-grs3-28), bt, hbl, cpx, chl ± sil grt-bt alteration with sil: grt (alm54-69-prp13-31-grs3-23), pl (mainly andesine, locally an), qtz, bt, sil grt-bt alteration without sil: (pl (xab0.56–0.57), grt (alm53-75-prp13-31-grs4-13), qtz, bt, chl) unaltered amphibolite, slightly retrogressed (hbl, pl (xab0.31-0.35),cpx, czo) 508405 508406 <2 ppb 23 ppb 508407 508408 508409 508410 508411 <2 ppb 12 ppb 4 ppb <2 ppb <2 ppb 508412 508413 508414 508415<2ppb <2ppb 508381 508382 508383 508384 508385 508386 508387 508388 508389 508390 508391 <2 ppb <2 ppb 3 ppb <2 ppb 46 ppb 34 ppb 20 ppb 38 ppb 672 ppb <2 ppb <2 ppb 508392 grt-bt-pl thermobarometry: 540°–620°c 4.5 ± 1 kbar profile aprofile b (cliff profile) 4 ppb <2 ppb 3 ppb 20 ppb <2 ppb 5 m petrographical sample (polished thin section) geochemical sample (fa gold)38 ppb 67 ppb sillimanite garnet sulphides qtz vein / qtz-blebs pegmatite silicified (schist) grt-rich schist bt-rich schist layered amphibolite (type 1) homogeneous amphibolite (type 2) lithogeochemical sample (whole-rock + trace elements + inaa gold) legend 51 a composition of alm53–69, prp21–31, grs3–12, depending on bulk composition. in some of the larger grains, the rims have slightly higher fe concentrations (alm62–72, prp20–25, grs3–10). metamorphic garnet in the unaltered wall rocks is unzoned and enriched in grossular (alm53–59, prp13–20, grs17–21). both types of garnet have very low ree contents (<2 ppm), and are depleted in lree. biotite in the alteration zone has aliv contents between 2.4 and 2.65 atoms per formula unit and mg/(fe+mg) ratios between 0.25 and 0.45. the composition of plagioclase from the amphibolites varies from andesine to anorthite; most grains can be classified as labradorite and bytownite. plagioclase in the hydrothermal alteration zone is depleted in ca and is mostly andesine. in order to achieve reliable pressure–temperature (p–t) estimates, only mineral cores of neighbouring minerals were used for geothermobarometry. due to the presence of retrograde reaction rims in some of the garnet grains, it was assumed that the mineral cores reflect the equilibrium mineral composition and were not altered by retrograde processes. p–t estimates on the alteration assemblage using the garnetbiotite-plagioclase-quartz geothermobarometer of wu et al. (2004) give conditions of c. 540–620°c and 4.5 ± 1 kbar (fig. 5). p–t pseudosection models using the computer program perplex developed by connolly (1990) confirm these conditions (koppelberg 2011). 500 µm hbhb grtgrt btbt btbt btbt silsil silsil silsil cr cu zn mo as au ni co sc la ce pr nd sm eu gd tb dy ho er tm yb lu –100 –50 0 50 100 150 200 –100 –50 0 50 100 150 200 –100 –50 0 50 100 150 200 a b c quartz-vein dominated alteration zone (vf: 1.14–1.62) silimanite-bearing alteration zone (vf: 0.5–0.85) g ain /lo ss (% ) g ain /lo ss (% ) g ain /lo ss (% ) sio2 al2o3 tio2 fe2o3 mgo mno cao na2o k2o fig. 3. photomicrograph illustrating the replacement of the regional metamorphic amphibolite facies mineral assemblages by hydrothermal garnet, biotite and sillimanite in sample ggu 508405. fig. 4. results of mass-balance calculations for quartz-vein dominated and sillimanite-bearing alteration zones. vf: volume factor – the change of volume of altered rock relative to unaltered rock. a: major elements, b: trace elements, c: rare-earth elements (ree). fig 5. thermobarometry results (see main text). the al2sio5 diagram is from holdaway & mukhopadhyay (1993). red: sample ggu 508384, green: 508385, blue: 508386, yellow: 508405. the al2sio5 triple point is at 500°c and 3.75 kbar. kyanite silimanite andalusite ky ky sil sil and and 200 300 500400 600 700 800 900 2 1 3 4 5 6 7 8 9 temperature (°c) pr es su re (k ba r) 5252 discussion and conclusion the majority of the world’s gold deposits formed in the archaean (c. 2.7 ga) as a result of crust-forming processes during collision events of converging continental plates (groves et al. 2005). these epigenetic deposits are called orogenic gold deposits, and occur in metamorphic terranes that mainly show greenschist facies metamorphism (groves et al. 1998). other orogenic deposits are known to have formed at amphibolite-facies metamorphic grades, and these are termed hypozonal deposits (groves et al. 1998). the replacement of regional, metamorphic, amphibolite-facies mineral assemblages by hydothermal minerals surrounding the goldbearing quartz veins as well as the late timing of quartz-vein formation by ductile, flexural slip folding (kolb et al. 2009), indicate that the gold mineralisation and associated hydrothermal alteration formed late in the metamorphic evolution on qilanngaarsuit. the low-pressure amphibolite-facies metamorphism in the surrounding amphibolites has been dated to c. 2715 ma (nutman & friend 2007), while the mineralisation probably occurred between 2660 and 2600 ma (kolb et al. 2013). the qilanngaarsuit mineralisation is, therefore, interpreted to represent a new example of hypozonal orogenic gold mineralisation in the godthåbsfjord gold province. the origin of other gold prospects (storø, qussuk) is still a matter of debate, and both metamorphosed epithermal and orogenic models have been proposed (garde et al. 2012; scherstén et al. 2012; kolb et al. 2013). the alteration systematics, timing and conditions of the gold mineralisation on qilanngaarsuit are, however, similar to those of other gold occurrences, in particular storø, in the godthåbsfjord gold province. the deposits are spatially closely associated with a major terrane boundary, the ivinnguit fault, suggesting that this shear zone may have acted as a major pathway for the gold-bearing fluids between c. 2660–2600 ma. acknowledgements the professor dr. karl-heinrich heitfeld-stiftung is thanked for financial support. the work benefited from valuable comments and discussions with susan giffin and nicolas stoltz. references allaart, j.h. 1982: geological maps of greenland 1:500 000. map sheet no. 2, frederikshåb isblink – søndre strømfjord. copenhagen: geological survey of denmark and greenland. chadwick, b. 1981: field relations, petrography and geochemistry of archaean amphibolite dykes and malene supracrustal amphibolites, northwest buksefjorden, southern west greenland. precambrian research 14, 221–259. connolly j.a.d. 1990: multivariable phase diagrams: an algorithm based on generalised thermodynamics. american journal of science 290, 666–718. garde, a.a., whitehouse, m. & christensen, r. 2012: mesoarchean epithermal gold mineralisation preserved at upper amphibolite-facies grade, qussuk, southern west greenland. economic geology 107, 881–908. gresens, p.l. 1967: composition-volume relationships of metasomatism. chemical geology 2, 47–65. groves, d.i., goldfarb, r.j., gebre-mariam, m., hagemann, s.g. & robert, f. 1998: orogenic gold deposits: a proposed classification in the context of their crustal distribution and relationship to other gold deposit types. ore geology reviews 13, 7–27. groves, d.i., condie, k.c., goldfarb, r.j., hronsky, j.m.a. & vielreicher, r.m. 2005: secular changes in global tectonic processes and their influence on the temporal distribution of gold-bearing mineral deposits. economic geology 100, 203–224. holdaway, m.j. & mukhopadhyay, b. 1993: a re-evaluation of the stability relations of andalusite: thermochemical data and phase diagram for the aluminum silicates. american mineralogist 78, 298–315. kolb, j., stensgaard, b.m., schlatter, d.m. & dziggel, a. 2009: controls of hydrothermal quartz vein mineralization and wall-rock alteration between ameralik and sermilik, southern west greenland. danmarks og grønlands geologiske undersøgelse rapport 2009/25, 76 pp. kolb, j., dziggel, a. & schlatter, d.m. 2013: gold occurrences of the archean north atlantic craton, southwestern greenland: a comprehensive genetic model. ore geology reviews 54, 29–58. koppelberg, m. 2011: geochemische und petrologische charakterisierung einer goldführenden hydrothermalen alterationszone auf der insel qilanngaarsuit, sw grönland, 70 pp. unpublished diploma thesis, institut of mineralogy and economic geology, rwth aachen university. nutman, a.p. & friend, c.r.l. 2007: adjacent terranes with c. 2715 and 2650 ma high-pressure metamorphic assemblages in the nuuk region of the north atlantic craton, southern west greenland: complexities of neoarchaean collisional orogeny. precambrian research 155, 159–203. scherstén, a., szilas, k., creaser, r.a., næraa, t., van gool, j.a.m. & østergaard, c. 2012: re-os and u-pb constraints on gold mineralisation events in the mesoto neoarchaean storø greenstone belt, storø, southern west greenland. precambrian research 200-203, 149-162. schlatter, d.m. 2009: petrographic and lithogeochemical surface data from the new gold occurrence on qilanngaarsuit island, southern west greenland. in: kolb, j & kokfelt, t.: annual workshop on the geology of southern west greenland realted to field work: abstract volume 1, geus rapport 2009/94, 18–21. wu, c.-m., zhang, j. & ren, l.-d. 2004: empirical garnet-biotite-plagioclase-quartz (gbpq) geobarometry in mediumto high-grade metapelites. journal of petrology 45, 1907–1921. authors’ addresses m.k., a.d., f.m.m., institute of mineralogy and economic geology, rwth aachen university, wüllnerstraße 2, 52056 aachen, germany; e-mail: koppelberg@iml.rwth-aachen.de d.m.s., helvetica exploration services gmbh, carl spitteler strasse 100, ch-8053 zürich, switzerland. j.k., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. geological survey of denmark and greenland bulletin 41, 2018, 33-38 33 during august 2017, as part of the habitat mapping of natura2000 areas, a geophysical survey of a large area within the skagerrak was undertaken by the geological survey of denmark and greenland. in this article, we use the acquired data to discuss the geology of tannis bugt (fig. 1), a large shallow bay at the north-west coast of vendsyssel. the bay extends 40 km between hirtshals in the west and skagen in the east forming the northern-most danish skagerrak coast. geological setting and glacial history in northern vendsyssel, the late quaternary succession is 250 m thick and underlain by cretaceous or triassic strata (håkansson & pedersen 1992). tertiary sediments are assumed to be absent (sandersen et al. 2009). vendsyssel was subjected to several ice advances during the saalian and weichselian, which resulted in a succession that is generally dominated by marine to glaciolacustrine clay (knudsen et al. 2009; larsen et al. 2009a). the oldest quaternary deposits in vendsyssel are represented by the clay-rich skærumhede till formation, which is followed by the marine lower skærumhede clay formation. the till units of the brønderslev and åsted formations of the early to middle weichselian reflect a change to a colder environment. boulders are found throughout the till units (pedersen 2005; knudsen et al. 2009; larsen et al. 2009a; larsen et al. 2009b). the changing climate resulted in ice retreat and the deposition of the marine sediments of the upper skærumhede clay formation, which gradually initial observations of the shallow geology in tannis bugt, skagerrak, denmark matthew j. owen, nicky h. witt, zyad al-hamdani, niels nørgaard-pedersen, katrine j. andresen and jørgen o. leth fig. 1. locality map. a: regional setting of the study area. jb: jammerbugt. rk: rubjerg knude. b: tannis bugt with location of data shown. 57°38´n 57°38´ 10°13´e 10°13´e 10°14´e 10°14´e camera view a24 sand bank area areas of ridged outcrop ns020 x l02 200 m a b 8 14 m fig. 2. a: seabed morphology with location of innomar lines xl02 and ns020 (shown in fig. 4). b: video view of the sea floor (a24). 100 km fig. 1b fig. 2 natura2000 survey area fig. 3 fig. 5 b a 6°e 60°n 57°n 6°e 12°e 60°n 10°e 10°30´e 57°39´n 5 km norwegian channel 12°e 57°44´n 10°30´e 57°39´n rk jb vendsyssel tannis bugt fig. 4d fig. 4a fig. 4cfig. 4b skagerrak north sea © 2018 geus. geological survey of denmark and greenland bulletin 41, 33–38. open access: www.geus.dk/bulletin http://www.geus.dk/bulletin 3434 developed into the glaciolacustrine lønstrup klint formation, and the transition continues into terrestrial and glacial sediments of the kattegat till formation. the late weichselian sediments consist of several phases of till and glaciolacustrine units. by c. 18 ka bp, as the ice retreated, large parts of vendsyssel were inundated by the sea leading to the deposition of the marine vendsyssel formation. associated with the advance and retreat of ice sheets, the area was strongly influenced by isostatic rebound with a regression occurring from the late glacial to the present day. raised late-glacial shorelines are found up to c. 60 m above sea level (richardt 1996). a number of glaciotectonic complexes have been mapped in northern jylland (fig. 1a) and tend to exhibit thrusting from the east or north-east (pedersen & boldreel 2015). the oceanographic setting has varied in response to the interplay between isostasy and eustasy. as such, the setting is inferred to have varied between an isolated embayment, extending from the atlantic via the norwegian channel, to the present-day setting as a deep-water component of the north sea. a. innomar line ns011, west b. innomar line ns011 east west east west east 10.3 km 9.8 km areas of deformation, with loss of internal structure area of deformation, with loss of internal structure fault with offset observed folding of reflectors lower unit consists of inclined, folded parallel and sub-parallel reflectors truncated at the unit’s upper boundary surface sand unit surface sand unit hummocky reflector with no signal penetration o verlaps w ith eastern section c. reflector geometry zonation 19.3 km west east west section part a east section part b dz1 dz2 dz3 dz4 dz5 deeper, more steeply inclined reflectors zone a zone b zone c zone c zone d zone e zone f zone g a b c d e ffd g crosses xl02 10 20 30 10 20 30 10 20 30 10 20 30 d ep th b el ow se a lev el (m ) d ep th b el ow se a lev el (m ) o verlaps w ith w estern section fig. 3. wsw–ene-orientated innomar profile ns011. a: western part. b: eastern part. c: full profile with zones described in table 1. dz: deformation zone. vertical exaggeration 70 times. 35 marine survey the dataset obtained in the survey consists of multibeam bathymetry, side-scan sonar and innomar sub-bottom profiler data. some 2150 km of data were acquired using the vessel r/v aurora (fig. 1b), covering the northern coast of vendsyssel between skagen and hirtshals. wsw–eneorientated nearshore lines were spaced at 300 m intervals and n–s-orientated crossing lines at 15 km intervals. in addition, 24 video views were made of the sea floor in the shallow-water area. camera view at site a24 – the site is located at a water depth of 12 m in an area characterised by a high-amplitude, irregular, acoustic seabed facies, in a depression adjacent to a low-amplitude shallow-banked area (fig. 2a). to the east an adjacent depression is characterised by a similarly high-amplitude acoustic facies, but traversed by nw–se-orientated linear ridges. when visible, the apparent strike of the bedding is generally se, though in some instances a ne strike is also visible. the camera view (fig. 2b) shows a mixture of coarse gravel, cobbles and boulders. the larger clasts appear to be sub-rounded to well rounded. innomar data – figures 3 to 5 show a selection of the innomar data from tannis bugt. the facies visible in these figures are representative of those observed across the site. the w–e-orientated innomar profile ns011 (fig. 3) shows a semi-transparent upper unit, which is intermittently present. it overlies a deeper unit with reflectors that dip to the north-east. an additional deeper unit, with more steeply inclined reflectors, is occasionally visible. to aid description the data shown in fig. 3 are divided into zones a to g from west to east (table 1). at least five zones of deformation (dz1–dz5; fig. 3) are apparent in the innomar data. indications of deformation include folding, offset reflectors and loss of structure. some of these complex features are shown in more detail in fig. 4. innomar examples (4a–c) show deformed reflectors, areas of lost structures and apparent faults. the side-scan example (fig. 4d) shows an area of wave-like bedding outcropping at the seabed. figure 5 shows the n–s-orientated innomar line xl02, which crosses the profile ns011 shown in fig. 3 as well as the area shown in fig. 2. as with the data shown in fig. 3, an upper semi-transparent unit rests unconformably on a unit with truncated dipping reflectors. in the southern half of the profile, the upper unit thickens, where it is associated with large sand banks. the reflectors of the lower unit show steeper inclination in the northern third of fig. 5, where inclination of 1.8° is found. near the crossing of ns011, shown in fig. 3, dz3 (fig. 3, table 1) is visible and marks a divide between the more steeply inclined reflectors farther north and the more gently inclined reflectors to the south, where inclinations of 0.7° are measured. at the far south, reflectors outcrop at the seabed with an inclination of 0.5°. blanking of the innomar data is particularly apparent in fig. 5 north of dz3; here reflectors are clearly cut by blanking zones rising from depth. south of dz3 penetration ends abruptly and is variable. some deformation is visible in fig. 5 (folded reflectors below the thickening surface unit and some loss of structure at dz3), though it is less apparent than in the data shown in figs 3, 4. table 1. summary of innomar data shown on line ns011 (fig. 3). zone length (m) facies characteristics representative reflector inclination a 3270 upper unit with a smooth and convex up upper boundary. the lower boundary shows undulation and a hummocky character. no innomar penetration below. na b 4650 upper unit thins and pinches out eastward. the seabed becomes increasingly incised with small notches moving east. this unit rests unconformably on a lower unit characterised by northeastwardly dipping reflectors. folded and offset reflectors are apparent (dz1 and dz2). 0.3° c 2640 an upper unit is generally present, though it pinches out occasionally and is <1 m thick. below, reflectors of the lower unit are inclined and appear to be more densely spaced, with deformation visible (dz3). the upper surface of the lower unit becomes more irregular to the east. 0.9° d 1800 upper unit is present, though <1 m thick, in the west and pinches out and is absent in the east. curvilinear, inclined reflectors characterise the lower unit, an apparent loss of internal structure is also visible (dz4). 0.6° e 930 lack of penetration, cuesta-like scarps of ~1 m height visible on the seabed. na f 4750 a semi-transparent upper unit with a lens-like morphology is intermittently present and <2 m thick. below, densely spaced reflectors are inclined giving the seabed a cuesta morphology where they outcrop. innomar penetration is variable. area of folding in the east (dz5). 0.9° fd deeper, more steeply inclined reflectors separate from the shallower units. 1.5° g 2070 intermittent upper unit, with a semi-transparent facies and lens-like morphology, resting unconformably atop of the irregular upper surface of the lower unit with inclined linear reflectors. the seabed shoals and innomar penetration reduces in the east. 1.0° 3636 discussion interpretation of innomar and video data – the upper units visible in the innomar data are interpreted as holocene marine sand and gravel. interpretation of the deeper unit is complicated by limited penetration from the innomar system and no ground control. however, as indicated by the zoning in figs 3 and 5, and the facies characteristics outlined in table 1, there is a clear variation in reflector geometry, and we make some provisional observations. one of the most striking aspects of the innomar data is the inclined nature of the reflectors, representing bedding planes, and the truncated, often irregular upper surface. as shown in table 1, inclination (as distinct from angle of dip due to the oblique orientation of the survey lines with respect to direction of dip) is low and measures between 0.3° and 1.5° within the w–e-orientated ns011 profile; it is generally greater in the east. measured inclinations from the n–sorientated xl02 profile (fig. 5) are greater, with 1.8°. there is also clear variation in angle between zones b–d, f and g (zones a and e do not show inclined reflectors). the parallel bedding appears to have undergone deformation (folds and faults are visible in figs 3, 4) in the centre and east of the site. some deformation is apparent in fig. 5, though it is less clear. areas of deformation separate zones b and c, c and d, and f and g. the reflective facies within the zones are distinct and display features indicative of further deformation. shown in fig. 5, there is a clear reduction in bedding inclination between zones c and b. innomar penetration is reduced by a factor of three in zone b compared to zone c and blanking cuts the bedding within zone c (indicating gas migration), whereas it does not in the more depth-limited imaging of zone b. this indicates a change in lithology between zones, with zone b potentially consisting of a less permeable and more dense material. referring to fig. 3, the reflectors in zone d display a lower angle of inclination and appear folded. zone e is characterised by no sub-surface penetration and the presence of cuesta-like scarps on the seabed. gas does not appear to be present in this location and the surface unit is thin or absent, therefore the lack of signal could be due to the presence of a massive or more dense unit. fig. 4. examples of deformations within tannis bugt. a–c: innomar data shown with a vertical exaggeration of 28 times. d: side-scan sonar data. d ep th b el ow se a lev el (m ) 20 30 a. deformation zone 3, profile ns10 b. deformation zone 4, profile ns11 d. outcropping fold, profile ns04 side-scan sonar 1.60 km 20 d ep th b el ow se a lev el (m ) 1.54 km 15 25 d ep th b el ow se a lev el (m ) c. deformation zone 5, profile ns12 0.77 km west east loss of structure 20 15 bending of reflectors fault irregular outcropping surface west eastupward bending reflectorsapparent fault facies boundary, associated with change in innomar penetration gas west eastloss of structure and deformed bedding chaotic reflections, loss of structure facies boundary gas curved, wave-like bedding outcropping at surface facies boundary linear bedding 50 km 57°42.3´n 57°42.24´n 10°21.84´e10°21.72´e 37 inclined reflectors are again apparent in zone f, sub-surface penetration varies and is generally limited. outcropping bedding is apparent in zone g, matching the acoustic facies visible on the side scan record: where outcropping strata are visible as ridges (fig. 2a). cobbles and boulders visible in camera view a24 (fig. 2b), where innomar data show high amplitude seabed in an area with outcropping strata, appear to have two possible origins: either a winnowed deposit from a quaternary till or drop-stone unit, or eroded pre-quaternary bedrock. the density of the clasts visible would require a large, though not impossible, initial thickness of till to yield such a high number of clasts. the presence of pre-quaternary geology would be unexpected; however, shown by the bedding visible in figs 3 and 5, outcrop in the location of camera site a24 is stratigraphically older than the lithological units located north and eastward, which extend for up to 20 km. origin and contributing geological processes – one of the more puzzling aspects of the geology of tannis bugt is the difference to the onshore data which show a succession of till, glaciomarine and lacustrine units that are incised by tunnel valleys covered by a thick succession of late glacial and holocene marine deposits (larsen et al. 2009a; sandersen et al. 2009). whilst the change in facies observed in the innomar data could occur due to transitions between till and glaciomarine or lacustrine units, the bedding inclination is not so simple to explain. it is possible that the deeper facies could represent crossbedding associated with the development of the skagen spit, or palaeo-skagen spit(s), with marine transgressive erosion surfaces noted on the north-western coastline of the present spit, which is migrating eastward (nielsen & johannessen 2009). however, the observed facies are too deep and, due to the regressive nature of the shoreline, would have been significantly deeper prior to the early holocene (richardt 1996). the inclination and north-eastward dipping of the beds in tannis bugt could indicate deposition within a prograding system during the development of the shallow-water area off vendsyssel, which extends some 20 km from the coastline. the truncated, inclined facies shown in figs 3, 5 are similar to a delta facies which may undergo deformation via gravity-driven processes (postma 1995; patruno et al. 2015). however, if a delta complex is present in tannis bugt, the age and sediment source are unknown. early quaternary deltaic units are present in the central and southern north sea (lamb et al. 2018), but these dip in a westerly direction. if deltaic, the ne–se-direction of dip observed in tannis bugt fig. 5. n–s-orientated innomar profile xl02. see fig. 2 for location. vertical exaggeration 40 times. innomar profile xl02 north south 10 20 30 40 crosses ns011 crosses ns020 zone c zone b 10 20 30 40 5.3 km folding of reflectors surface sand unit gas blanking area of deformation dz3 high amplitude ‘bright spots’ ? ? d ep th b el ow se a lev el (m ) 3838 would indicate progradation into the skagerrak and norwegian channel. this suggests that the inclined bedding could represent subaqueous shelf-prism clinoforms (patruno et al. 2015), deposited within the skagerrak embayment during low stands (producing the observed sw–ne-orientated strike) and as a consequence of the stronger south-west flow from the north sea during high stands (producing the sw– ne-orientated strike). glaciotectonic activity is documented near tannis bugt, with complexes at rubjerg knude (pedersen 2005) and in jammerbugt (pedersen & boldreel 2017; fig. 1), and could also cause the observed facies. tannis bugt has a dip geometry (to the north-east) that is similar to jammerbugt, though the measured reflector inclination of c. 1–2° is less than the c. 6° thrust plane dip seen in the proximal part of the jammerbugt complex (pedersen & boldreel 2017). the lower angle of inclination could be due to limited penetration of the innomar data not revealing the unit’s true geometry, or it could represent a low-angled thrust component, such as a flat (pedersen & boldreel 2015). however, it is perhaps more likely that the facies are a result of the combined processes of clinoform deposition, causing the observed inclination, and glaciotectonic activity causing the deformation. if the deformation in tannis bugt is glaciotectonic, it is unclear whether it is part of a larger tannis–jammerbugt complex, or whether there are two separate complexes. in any case, their geographical proximity and similar direction of dip indicate that they could be formed by a similar process such as an advance from the norwegian channel ice stream. conclusion inclined and deformed bedding planes dipping predominantly to the north-east have been identified in tannis bugt, off the north coast of vendsyssel. a bouldery area, representing either winnowed glacial unit or eroded pre-quaternary bedrock, has been identified near the coastline in the centre of the bay. if the inclined bedding represents a prograding sequence with an eroded upper surface then the outcrop long section could represent a unit of significant age, with the possibility of pre-quaternary geology exposed at the seabed. if the inclination and deformation are the result of glaciotectonic deformation then the site represents a newly discovered glaciotectonic complex. deeper penetrating multichannel seismic data and geological samples will be required to fully understand the geology of tannis bugt and how it relates to the development of northern denmark. acknowledgements we thank the danish environmental protection agency for permission to publish data acquired during the survey of natura2000 areas, the captain and crew of the r/v aurora for their assistance during the survey and orbicon a/s for the acquisition of video data. references håkansson, e. & pedersen, s.a.s. 1992: geologisk kort over den danske underground, map sheet. copenhagen: varv. knudsen, k.l., kristensen, p. & larsen, n.k. 2009: marine glacial and interglacial stratigraphy in vendsyssel, northern denmark: foraminifera and stable isotopes. boreas 38, 787–810. lamb, r.m., harding, r., huuse, m., stewart, m. & brocklehurst, s.h. 2018: the early quaternary north sea basin. journal of the geological society 175, 275–290. larsen, n.k., knudsen, k.l., krohn, c.f., kronborg, c., murray, a.s. & nielsen, o.l.e.b. 2009a: late quaternary ice sheet, lake and sea history of southwest scandinavia – a synthesis. boreas 38, 732–761. larsen, n.k., krohn, c.f., kronborg, c., nielsen, o.b. & knudsen, k.l. 2009b: lithostratigraphy of the late saalian to middle weichselian skaerumhede group in vendsyssel, northern denmark. boreas 38, 762–786. nielsen, l.h. & johannessen, p.n. 2009: facies architecture and depositional processes of the holocene–recent accretionary forced regressive skagen spit system, denmark. sedimentology 56, 935–968. patruno, s., hampson, g.j. & jackson, c.a.-l. 2015: quantitative characterisation of deltaic and subaqueous clinoforms. earth science reviews 142, 79–119. pedersen, s.a.s. 2005: structural analysis of the rubjerg knude glaciotectonic complex, vendsyssel, northern denmark. geological survey of denmark and greenland bulletin 8, 192 pp. pedersen, s.a.s. & boldreel, l.o. 2015: thrust-fault architecture of glaciotectonic complexes in denmark. geological survey of denmark and greenland bulletin 33, 17–20. pedersen, s.a.s. & boldreel, l.o. 2017: glaciotectonic deformations in the jammerbugt and glaciodynamic development in the eastern north sea. journal of quaternary science 32, 183–195. postma, g. 1995: sea-level-related architectural trends in coarse-grained delta complexes. sedimentary geology 98, 3–12. richardt, n. 1996: sedimentological examination of the late weichelian sea-level history following deglaciation of northern denamrk. in: andrews, j.t. et al. (eds): late quaternary palaeoceanography of the north atlantic margins. geological society (london), special publication 111, 261–273. sandersen, p.b.e., jørgensen, f., larsen, n.k., westergaard, j.h. & auken, e. 2009: rapid tunnel-valley formation beneath the receding late weichselian ice sheet in vendsyssel, denmark. boreas 38, 834–851. authors’ addresses m.j.o., n.h.w., z.a.-h., n.n.-p., j.o.l., geological survey of denmark and greenland, c.f. møllers allé 8, dk-8000 aarhus c, denmark. e-mail: mow@geus.dk. k.j.a., department of geoscience, aarhus university, høegh-guldbergs gade 2, dk-8000 aarhus c, denmark. mailto:mow@geus.dk geological survey of denmark and greenland bulletin 31, 2014, 83-86 83 katabatic winds and piteraq storms: observations from the greenland ice sheet dirk van as, robert s. fausto, konrad steffen and the promice project team* * andreas p. ahlstrøm, signe b. andersen, morten l. andersen, jason e. box, charalampos charalampidis, michele citterio, william t. colgan, karen edelvang, signe h. larsen, søren nielsen, martin veicherts and anker weidick in 2007 the programme for monitoring the greenland ice sheet (promice) was initiated to observe and gain insight into the mass budget of greenland ice masses. by means of in situ observations and remote sensing, promice assesses how much mass is gained as snow accumulation on the surface versus how much is lost by iceberg calving and surface ablation (ahlstrøm et al. 2008). a key element of promice is a network of automatic weather stations (awss) designed to quantify components of the surface mass balance, including the energy exchanges contributing to surface ablation (van as et al. 2013). the use of these aws observations is not limited to studies of ice-sheet mass balance. promice contributes to cryo net (www.globalcryospherewatch.org/cryonet), the core net work of surface measurement sites of the world meteorological organization (wmo) global cryosphere watch. by real-time delivery through wmo, promice observations contribute to improve both operational forecasting and climate analysis in the data-sparse arctic. the greenlandic population, highly dependent on accurate forecasting of weather conditions, benefits directly from these real-time observations. for instance, extreme surface wind speeds are a high-risk element in greenland. the third-highest wind speed observed at the surface of the earth (93 m/s or 333 km/h), was recorded in a 8–9 march 1972 storm at thule in north-west greenland (stansfield 1972). in this paper, we discuss the extent to which the greenland ice sheet generates its own near-surface wind field. we use promice data to gain insight into the interaction between air temperature, radiation and gravity-driven katabatic winds. we focus on a particularly powerful spring storm in 2013 that contributed to a fatality on an ice-sheet ski traverse attempt (linden 2013). weather stations on the greenland ice sheet the original promice network consisted of fourteen awss in seven ablation regions of the greenland ice sheet, with each region monitored by a lower (l) and an upper (u) elevation station (fig. 1; ahlstrøm et al. 2008). promice has collaborated logistically and financially with other projects in the regions of the tas, qas, nuk and kan stations, leading to the installation of eight additional awss. the promice © 2014 geus. geological survey of denmark and greenland bulletin 31, 83–86. open access: www.geus.dk/publications/bull fig 1. map of greenland showing the locations of automatic weather stations on the ice sheet and on local ice caps. 250 km zak qas nuk kan mal thu upe kpc sco mit tas fig. 522 50 greenland climate network promice network other geus stations (co)funded by other projects 2250 3000 2750 2500 1000 1500 1750 20 00 25 00 22 50 17 50 1500 12 50 2000 8484 aws sites have been selected to complement the greenland climate network (gc-net), which chiefly monitors the icesheet accumulation area (steffen et al. 1996). continuous promice aws observations include: air temperature (c. 2.7 m above surface), barometric pressure, air humidity, wind speed and direction (c. 3.1 m above surface) as well as downand upward solar (shortwave) and terrestrial (longwave) radiation. the awss also record temperature profiles in the upper 10 m of the ice, gps-derived location and diagnostic parameters such as station tilt angles. a pressure transducer and two sonic rangers measure snow and icesurface height change associated with ablation and accumulation (fausto et al. 2012). all data and metadata including sensor specifications are available at www.promice.org. here, we use averaged values of air temperature, wind speed and direction, and radiation components. single wind measurements have an uncertainty of 0.3 m/s and 3° (van as 2011) and are not adjusted for shifts in tilt, rotation or measurement height as this does not impact the outcome of this study. we also combine gc-net and promice temperature data to give the most complete observed meteorological depiction of the greenland ice sheet currently possible. we calculated the daily average near-surface air temperature across the ice sheet between 2008 and 2013 by means of inverse-distance interpolation between as many as 32 awss that operated on a given day. we also determined the daily average vertical near-surface air-temperature lapse rate by means of a linear least-squares fit to all available data. atmospheric temperature and stability the average near-surface air temperature over the greenland ice sheet has a distinct annual cycle with minimum (winter) values between –20°c and –40°c (fig. 2a). during the relatively short summer, temperatures are often around –5°c and are less variable due to (1) reduced cyclonic activity and (2) surface melting over large parts of the ice sheet. the latter is a moderating factor because near-surface temperatures are limited to near freezing. since 2008, ice-sheet average air temperatures above 0°c have only been recorded on five days (11–13 and 28–29 july 2012) during which surface melting occurred over nearly the entire ice sheet (e.g. nghiem et al. 2012). the average near-surface air-temperature lapse rate over the ice sheet exhibits a reversed cyclicity as compared to air temperature with winter values often exceeding a 10°c decrease per vertical kilometre (fig. 2b). assuming a textbook value of a 6.5°c/km free-atmospheric lapse rate to be representative year 20092008 2010 2011 2012 20092008 2010 2011 2012 a b a ir te m pe ra tu re (° c ) 0 –20 –10 –30 –40 te m pe ra tu re la ps e ra te (° c /k m ) 15 10 5 net surface radiation (w/m2) 0–50 50 100 150 200 1 month of year 2 3 4 9 10 11 125 6 7 8 a b w in d sp ee d (m /s ) w in d sp ee d (m /s ) 8 9 6 7 4 5 3 2 10 6 2 8 4 scokpc tas qas nuk kan upe thu fig. 2. a: daily average (black) and 31-day average (red) air temperature over the greenland ice sheet as determined from interpolated weather station observations from the gc-net and promice network. b: same, but vertical near-surface temperature lapse rates. the dashed line shows a lapse rate of 6.5°c/km above which air masses are increasingly unstable. fig. 3. a: the average annual cycle in wind speed at the promice sites. lines are drawn for each weather station, but only if three years of good data are available. b: the monthly average wind speed versus the net (shortwave + longwave) radiation budget. for locations of the stations see fig. 1. 85 of the threshold between stable and unstable conditions over greenland, this suggests that the near-surface atmosphere is commonly less buoyant (denser) at higher elevations than air at lower elevations. in a free atmosphere such a density difference over a few vertical kilometres would trigger an immediate adjustment through convection. over the large horizontal scale of the greenland ice sheet, the actual density gradients are roughly two orders of magnitude smaller, which adds insignificantly to the force balance. figure 2b illustrates that during winter, the high elevation interior of the ice sheet cools more than lower elevation regions near the margin. as a result, the shallow (c. 100 m thick) stable atmospheric boundary layer that blankets the ice sheet attains an even larger temperature deficit compared to the free atmosphere at high elevation in winter. the larger this temperature deficit relative to the free atmosphere, the larger the density difference relative to the free atmosphere, and thus the larger the gravitational acceleration of the shallow boundary layer. this katabatic force increases linearly with increasing surface slope. katabatic winds winds over the greenland ice sheet are strongest in winter (e.g. steffen & box 2001), as observed at every promice aws (fig. 3a). while part of this increase is due to lower wintertime pressure and more frequent passage of cyclonic systems, the primary cause of stronger winter winds is surface radiative cooling. this well-known forcing mechanism of katabatic wind is apparent from stronger winds at more negative surface net radiation (fig. 3b) and the strong correlation between the directions in slope and wind (see below). a negative radiation budget is common during winter due to little or no solar radiation at high latitudes when the upward emission of long-wave terrestrial radiation exceeds downward atmospheric radiation at the surface. the wind regimes over the ice sheet do differ between regions. winds are stronger at the higher-elevation awss due to the larger radiative cooling of the surface (provided a surface slope is present). the highest monthly-mean wind speed values in fig. 3b were recorded at kan_m and kan_u (1270 and 1840 m a.s.l., red), and tas_u and tas_a (570 and 900 m a.s.l., blue). piteraq storms the wind regimes at kan and tas are shown in a case study of the 2012/2013 winter (fig. 4). figure 4a illustrates that low-wind winter conditions are rare at kan_u, promice station at highest elevation. figure 4b shows the dominant katabatic nature of winter winds. nearly all measurements from kan_u show the wind to blow from upslope direction (c. 90°, east), albeit deflected to the right (c. 135°, south-east) by the coriolis effect due to the earth’s rotation. typically, wind speeds at tas_u are lower (but still nonzero) due to the weaker radiative cooling at lower elevation. katabatic forcing also dominates here, given the persistent non-zero winds originating from the upslope direction of c. 0° (north) and more westerly directions due to coriolis forcing. the major difference between the two data series in fig. 4 is the frequency of strong wind events exceeding c. 20 m/s, which are more common in the tas region. in the strongest storms, the wind direction pivots towards the regional freeatmospheric flow (fig. 4b). these storms are known in greenland as piteraqs, and build up momentum due to the alignment of katabatic and large-scale (geostrophic) forcing (oltmanns et al. 2014). these notorious storms have repeatedly caused severe damage to the towns such as tasiilaq. the piteraq on 27 april 2013 (fig. 4a), which jeopardised a sport expedition on the ice sheet (linden 2013), was exceptionally strong at tas_u in the context of the 2008 to 2013 promice observational period, with 10-minute average wind speeds exceeding 42 m/s (150 km/h). during this event, four persons (c. charalampidis, w.t. colgan, h. machguth and d. van as) wind direction (°) 900 180 270 360 jan month of 2013 feb mar mayapr a b w in d sp ee d (m /s ) w in d sp ee d (m /s ) 40 30 10 20 0 40 20 0 30 10 east south west tas_u kan_u 27 april 2013 fig. 4. a: hourly average wind speed at tas_u and kan_u weather stations. the piteraq on 27 april 2013 is clearly visible in the tas_u observations. b: same, but wind speed plotted versus wind direction for the period from october 2012 to may 2013. 8686 from the geological survey of denmark and greenland were in the field at kan_u, and although they experienced wind speeds approximately one third of those at tas_u (c. 300  km to the east) the white-out and heavy snowdrift yielded conditions too dangerous for them to leave shelter. satellite images from the 2013 piteraq event show that a large region was affected (fig. 5). the striping on the ice sheet in the top left corner of the lower image shows the wind direction with snow transported toward and past the ice sheet margin. large areas of sea and fjord ice disintegrated, and the 5–13 km wide sermilik fjord, into which helheimgletscher calves, was cleared of ice. clearly, katabatic winds and especially the piteraqs, have a large impact on the ice sheet and its immediate surroundings. given increasing commercial activity around the periphery of the greenland ice sheet, there is a growing impetus for understanding these winds and their response to climate change. regional atmospheric model projections until the year 2100 suggests that while climate change will likely result in weaker winds in greenland’s flat interior, stronger winds may occur in steeper regions around the ice sheet periphery (gorter et al. 2013). acknowledgements promice is funded by the danish ministry of climate, energy and building and is operated by the geological survey of denmark and greenland. several weather stations are (co)funded by the greenland analogue project, the refreeze project and the greenland climate research centre. references ahlstrøm, a.p. & the promice project team 2008: a new programme for monitoring the mass loss of the greenland ice sheet. geological survey of denmark and greenland bulletin 15, 61–64. fausto, r.s., van as, d., ahlstrøm, a.p. & citterio, m. 2012: assessing the accuracy of greenland ice sheet surface ablation measurements by pressure transducer. journal of glaciology 58(212), 1144–1150. gorter, w., van angelen, j.h., lenaerts, j.t.m. & van den broeke, m.r. 2013: present and future near-surface wind climate of greenland from high resolution regional climate modelling. climate dynamics 42, 1595–1611. linden, m. 2013: greenland death: how 100 mph ice storms can blow in without warning. the independent, 1 may 2013. london, uk. nghiem, s.v., hall, d.k., mote, t.l., tedesco, m., albert, m.r., keegan, k., shuman, c.a., digirolamo, n.e. & neuman, g. 2012: the extreme melt across the greenland ice sheet in 2012. geophysical research letters 39, l20502. oltmanns, m., straneo, f., moore, g.w.k. & mernild, s.h. 2014: strong downslope wind events in ammassalik, southeast greenland. journal of climate 27, 977–993. stansfield, j. 1972: the severe arctic storm of 8–9 march 1972 at thule air force base, greenland. weatherwise 25, 228–233. steffen, k. & box, j.e. 2001: surface climatology of the greenland ice sheet: greenland climate network 1995–1999. journal of geophysical research: atmospheres 106(d24), 33951–33964. steffen, k., box, j.e. & abdalati, w. 1996: greenland climate network: gc-net. in: colbeck, s.c. (ed.): glaciers, ice sheets and volcanoes: a tribute to mark f. meier. crrel special report 96-27, 98–103. van as, d. 2011: warming, glacier melt and surface energy budget from weather station observations in the melville bay region of northwest greenland. journal of glaciology 57(202), 208–220. van as, d., fausto, r.s., colgan, w.t., box, j.e. & promice project team 2013: darkening of the greenland ice sheet due to the melt-albedo feedback observed at promice weather stations. geological survey of denmark and greenland bulletin 28, 69–72. 24 april 2013 tas_u mit tasiilaq 27 april 2013 sermilik helheimgletscher authors’ addresses d.v.a. and others, except k.s., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: dva@geus.dk k.s., swiss federal institute for forest, snow and landscape research (wsl), zürcherstrasse 111, ch-8903 birmensdorf, switzerland. fig. 5. modis satellite images of the tasiilaq region of south-east greenland on 24 and 27 april 2013, before and during a strong piteraq event. for location, see fig. 1. geological survey of denmark and greenland bulletin 4, 2003, pp 69-72 69 southern west greenland hosts a province of ultramafic alkaline rocks, including swarms of dykes traditionally described as kimberlites and lamproites (larsen 1991; jensen et al. 2002). since the mid-1990s, commercial diamond exploration has been focused on the sarfartoq region and the region south-east of maniitsoq (fig. 1), and has resulted in numerous reports of diamond-favourable indicator minerals from till sampling, finds of kimberlitic dykes, and recovery of diamonds from kimberlitic rocks. a new digital compilation of company data released from confidential status (jensen et al. 2003a) presents a comprehensive overview of exploration activities and results that have emerged since the survey’s first compilation of occurrences of kimberlitic and related rocks (larsen 1991). the new compilation in a gis (geographic information system) environment allows for refined assessment of the distribution, structural control and possible spatial and petrogenetic relationships that characterise the kimberlitic occurrences. in 2003, the geological survey of denmark and greenland (geus) and the government of greenland’s bureau of minerals and petroleum (bmp) went further than has been customary in investigating the economic potential of specific sites. four areas were temporarily closed to application for exploration licences, pending sampling and testing for diamond content of large samples of more than one tonne each from significant kimberlitic occurrences. additional characterisation and research initiated on these and other occurrences include magnetic mapping, detailed petrography and studies of mantle xenoliths, as well as indicator mineral chemistry. an extensive programme to determine the ages of kimberlitic and related rocks was also initiated in 2003. distribution of kimberlitic occurrences alkaline ultramafic dykes in the sisimiut–kangerlussuaq and sarfartoq regions intrude the border zone between the archaean craton and the palaeoproterozoic nagssugtoqidian orogen (fig. 1; secher & larsen 1980). the occurrences south-east of maniitsoq mark the southernmost extent of the alkaline province. the alkaline rocks of interest here have previously been described as kimberlites and lamproites (larsen 1991; jensen et al. 2002). this classification was questioned by mitchell et al. (1999), who consider that they are not typical kimberlites, but unusual ultramafic lamprophyres in that they are sometimes diamondiferous. in the absence of petrographic data for many of the occurrences the broader terms ‘kimberlitic’ and ‘lamproitic’ are applied here. three clusters of dykes have been recognised within the province during the last 20–30 years (larsen 1980, 1991; scott 1981). the ‘sisimiut cluster’, consisting mainly of 1214–1240 ma lamproitic and c. 590 ma kimberlitic dykes (larsen & rex 1992), is characterised by vertical e–w to se–nw trends. the ‘sarfartoq cluster’ has been described as a cone-sheet structure centred on the 600 ma sarfartoq cargeological survey of denmark and greenland bulletin 4, 69–72 (2004) © geus, 2004 investigating the diamond potential of southern west greenland sven monrad jensen and karsten secher fig. 1. map of kimberlitic and diamond occurrences of the west greenland alkaline province. framed areas enclose the three kimberlitic occurrences sampled for determination of diamond content in 2003. bonatite complex (larsen 1980). the 600 ma kimberlitic dykes of the ‘maniitsoq cluster’ have more variable orientations. the cone-sheet model for kimberlitic dykes around the sarfartoq carbonatite complex was largely based on dyke orientations in a broad e–w-trending valley transecting the core of the complex. with new knowledge of hundreds of additional kimberlitic occurrences in the region presented in jensen et al. (2003c), it now appears that other structural elements have controlled the emplacement of dykes. for example, some kimberlitic dykes follow the trends of the palaeoproterozoic kangâmiut dolerite dykes, as outlined by escher et al. (1970, 1976), in reworked as well as unreworked parts of the archaean basement. another example is an apparent predominance of n–s-trending kimberlitic dykes in a zone reaching far beyond the sarfartoq carbonatite complex. information from magnetic field data lends support to the hypothesis that kimberlitic dyke emplacement may be controlled by structures of regional character (jensen et al. 2003b, c). digital compilation of data exploration companies have produced a large volume of data relevant to diamond exploration, especially since 1994. the data include analyses of heavy minerals from till and stream sediment samples, dyke and boulder distribution maps, airborne and ground geophysical surveys, results of testing for diamond content of mini-bulk sampled dykes, drill logs, etc. a gis compilation of company exploration data now in the public domain (jensen et al. 2003a) constitutes the most extensive overview of kimberlitic rocks and diamond occurrences in greenland to date. the compilation contains scanned text and maps, and selected digital data from company assessment reports submitted to bmp in fulfilment of the standard terms for mineral exploration licences. a total of 146 company reports of relevance to diamond exploration, with a total of approximately 9250 pages of text, tables and maps, are included as pdf files. more than 53 000 tabulated analyses of indicator minerals from till and stream sediment are available, as are the details and results of drilling campaigns and tests for diamond content. 70 fig. 2. outcrop of one of the three kimberlitic occurrences from which more than 1000 kg of rock was collected (area 1 in fig. 1). width of dyke approx. 1 m. diamond occurrences and indicator mineral chemistry most of the approximately 600 diamonds reported to date in greenland are from just two areas, both located in the unreworked archaean craton (jensen et al. 2003a). some 95% of the stones are classified as microdiamonds, i.e. smaller than the minimum size recoverable in operating mine plants – typically those passing a 0.5 mm screen. another 20 microdiamonds from kimberlitic occurrences to the west of the sarfartoq carbonatite complex were reported in late 2003 (tuer 2003). the largest diamond from an in situ dyke reported to date is c. 1.7 mm in its longest dimension and has a weight of around 0.001 carat (1 carat = 0.2 g). although microdiamonds do not constitute an economic resource, they are important to the characterisation of kimberlites and evaluation of diamond deposits (rombouts 1995). nearly all of the reported diamonds have been recovered using caustic dissolution, a method that is usually adjusted to recover stones of all sizes down to around 0.1 or 0.15 mm. in addition, a few diamonds have been found in large stream sediment samples. some occurrences have also been subjected to dense media separation tests for larger diamonds, where up to 11 tonnes of kimberlitic rock have been processed (boucher 2000), but none of these tests have recovered any diamonds. all in situ diamond occurrences lie within areas outlined by diamond-favourable indicator minerals from till and stream sediment samples. on a local scale, however, kimberlite tracing using indicator minerals from till samples is not straightforward, probably due to the influence of complex glacial dynamics on the formation of the till deposits. the most diamond-favourable indicator minerals are distributed far beyond the areas with known diamonds. this observation, together with the postulated regional structural controls, suggests that the potential for diamonds is not restricted to the known occurrences. the potential appears to exist on both sides of the boundary between reworked and unreworked archaean basement. testing of three kimberlitic occurrences for diamond content in 2003, geus and bmp undertook sampling of three large occurrences of kimberlitic rocks for subsequent testing for diamond content using caustic dissolution. composite samples of approximately 1000 kg from each occurrence have been processed and examined for diamonds by a certified canadian testing laboratory. two of the occurrences are vertical dykes with a length of approximately 2500 m and a width of up to 2 m (figs 2–4). the third occurrence is a shallow-dipping sill with a length of at least 500 m and a thickness of 1–2 m. the sill and one of the long dykes lie well within the unreworked archaean craton, while the second long dyke lies a few kilometres inside the palaeoproterozoic deformed region (fig. 1). the test resulted in 125 diamonds recovered from the sampled dyke in area 1 (fig. 1), two diamonds from the dyke in area 2 and one diamond from the dyke in area 3 (jensen et al. 2004). the largest diamond recovered measures 0.74 × 0.63 × 0.54 mm, and the total weight of the 128 stones is 0.016 carat. 71 fig. 3. eclogitic xenolith in one of the dykes sampled for testing for diamond content (area 1 in fig. 1). scale bar is 2 cm. fig. 4. boulders of a 1.5 m wide and 2500 m long dyke sampled for testing for diamond content (area 3 in fig. 1). 72 indicator minerals from the same three kimberlitic occurrences have been separated, picked and analysed by electron microprobe. the indicator minerals studied are sub-calcic pyrope, eclogitic garnet, chrome-diopside, chromite, ilmenite and olivine. the indicator mineral chemistry will be used in conjunction with the diamond determination results to assess the diamond potential of the occurrences. the mineral chemistry fingerprint of the in situ occurrences may have important implications for the interpretation of existing indicator mineral chemistry data from till and stream sediment samples. research in progress the field work in 2003 included detailed magnetic mapping of the three occurrences using a proton magnetometer. interpretation of the geophysical field data is ongoing. petrographic and geochemical characterisation of the kimberlitic occurrences has been largely neglected in previous investigations, and accordingly a comprehensive programme to systematically study and classify the groundmass, mantle xenoliths and indicator minerals of these rocks has been initiated. an extensive programme to determine the age of these and many additional kimberlitic occurrences using the very precise u-pb in perovskite method (e.g. heaman et al. 2003) has been launched. some phlogopite-rich rocks will be dated by the rb-sr method. a total of around 35 age determinations have been commissioned. an updated version of the digital data compilation (jensen et al. 2003a) planned for 2004 will include the results of the ongoing survey testing and analytical work, as well as recently released company data, amounting to 3300 pages of text, tables and maps, 50 000 indicator mineral analyses, around 100 previously undescribed kimberlitic occurrences and a large volume of airborne geophysical data. acknowledgement the work reported on here has been supported and financed in part by the bureau of minerals and petroleum in nuuk. references boucher, d.r. 2000: 1999 assessment work report on the mini-bulk sampling programme, sarfartoq exploration licence, kalaallit nunaat, 18 pp. unpublished report, monopros ltd. for dia met minerals ltd., kelowna, b.c., canada (in archives of the geological survey of denmark and greenland, geus report file 21742). escher, a., escher, j.c. & watterson, j. 1970: the nagssugtoqidian boundary and the deformation of the kângamiut dyke swarm in the søndre strømfjord area. rapport grønlands geologiske undersøgelse 28, 21–23. escher, a., sørensen, h. & zeck, h.p. 1976: nagssugtoqidian mobile belt in west greenland. in: escher, a. & watt, w.s. (eds): geology of greenland, 76–103. copenhagen: geological survey of greenland. heaman, l.m., kjarsgaard, b.a. & creaser, r.a. 2003: the timing of kimberlite magmatism in north america: implications for global kimberlite genesis and diamond exploration. lithos 71, 153–184. jensen, s.m., hansen, h., secher, k., steenfelt, a., schjøth, f. & rasmussen, t.m. 2002: kimberlites and other ultramafic alkaline rocks in the sisimiut–kangerlussuaq region, southern west greenland. geology of greenland survey bulletin 191, 57–66. jensen, s.m., lind, m., rasmussen, t.m., schjøth, f. & secher, k. 2003a: diamond exploration data from west greenland. danmarks og grønlands geologiske undersøgelse rapport 2003/21, 50 pp. + 1 dvd. jensen, s.m., secher, k., rasmussen, t.m., tukiainen, t., krebs, j.d. & schjøth, f. 2003b: distribution and magnetic signatures of kimberlitic rocks in the sarfartoq region, southern west greenland. 8th international kimberlite conference, victoria, b.c., canada. extended abstracts cd-rom, 5 pp. jensen, s.m., secher, k., rasmussen, t.m., tukiainen, t., krebs, j.d. & schjøth, f. 2003c: distribution and magnetic signatures of kimberlitic rocks in the sarfartoq region, southern west greenland. 8th international kimberlite conference, victoria, b.c., canada. poster presentation [available as pdf on cd-rom from authors]. jensen, s.m., secher, k. & rasmussen, t.m. 2004: diamond content of three kimberlitic occurrences in southern west greenland. diamond identification results, field description and magnetic profiling. danmarks og grønlands geologiske undersøgelse rapport 2004/19, 41 pp. larsen, l.m. 1980: lamprophyric and kimberlitic dykes associated with the sarfartôq carbonatite complex, southern west greenland. rapport grønlands geologiske undersøgelse 100, 65–69. larsen, l.m. 1991: occurrences of kimberlite, lamproite and ultramafic lamprophyre in greenland. open file series grønlands geologiske undersøgelse 91/2, 36 pp. larsen, l.m. & rex, d.c. 1992: a review of the 2500 ma span of alkalineultramafic, potassic and carbonatitic magmatism in west greenland. lithos 28, 367–402. mitchell, r.h., scott smith, b.h. & larsen, l.m. 1999: mineralogy of ultramafic dikes from the sarfartoq, sisimiut and maniitsoq areas, west greenland. in: gurney, j.j. et al. (eds): proceedings of the viith international kimberlite conference 2, 574–583. cape town: red roof design cc. rombouts, l. 1995: sampling and statistical evaluation of diamond deposits. in: griffin, w.l. (ed.): diamond exploration into the 21st century. journal of geochemical exploration 53(1–3), 351–367. scott, b.h. 1981: kimberlite and lamproite dykes from holsteinsborg, west greenland. meddelelser om grønland, geoscience 4, 24 pp. secher, k. & larsen, l.m. 1980: geology and mineralogy of the sarfartôq carbonatite complex, southern west greenland. lithos 13, 199–212. tuer, j. 2003: hudson announces diamond results for west greenland program. news release nr2003-5, 1 p. vancouver, canada: hudson resources inc. (issued 1 october 2003). authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: smj@geus.dk << /ascii85encodepages false /allowtransparency false /autopositionepsfiles true /autorotatepages /all /binding /left /calgrayprofile (dot gain 20%) /calrgbprofile (srgb iec61966-2.1) /calcmykprofile (u.s. web coated \050swop\051 v2) /srgbprofile (srgb iec61966-2.1) /cannotembedfontpolicy /warning /compatibilitylevel 1.4 /compressobjects /tags /compresspages true /convertimagestoindexed true /passthroughjpegimages true /createjdffile false /createjobticket false /defaultrenderingintent /default /detectblends true /colorconversionstrategy /leavecolorunchanged /dothumbnails false /embedallfonts true /embedjoboptions true /dscreportinglevel 0 /syntheticboldness 1.00 /emitdscwarnings false /endpage -1 /imagememory 1048576 /lockdistillerparams false /maxsubsetpct 100 /optimize true /opm 1 /parsedsccomments true 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<feff4f7f752890194e9b8a2d5b9a5efa7acb76840020005000440046002065874ef65305542b8f039ad876845f7150cf89e367905ea6ff0c4fbf65bc63d066075217537054c18cea3002005000440046002065874ef653ef4ee54f7f75280020004100630072006f0062006100740020548c002000520065006100640065007200200035002e0030002053ca66f465b07248672c4f86958b555f3002> >> >> setdistillerparams << /hwresolution [2400 2400] /pagesize [595.000 842.000] >> setpagedevice geological survey of denmark and greenland bulletin 41, 2018, 9-12 9 oil and gas production from siliciclastic reservoirs has hitherto been in the danish central graben mostly from palaeogene and middle jurassic sandstone. the ravn field was the first upper jurassic field to start operation. the reservoir is composed of sandstone of the heno formation. production takes place at a depth of 4000 m, which makes ravn the deepest producing field in the danish north sea. the heno formation mainly consists of marine shoreface deposits, where foreshore, middle and lower shoreface sandstones constitute the primary reservoir. the results of this study of the diagenetic impact on the mineralogical composition, porosity and permeability are presented here. microcrystalline quartz has preserved porosity in the sandstone, whereas illite, quartz overgrowth and carbonate cement have reduced both porosity and permeability. geological background the ravn member of the heno formation is located on the heno plateau in the danish central graben (fig. 1; johannessen 2010). the ravn field was discovered in the ravn-1 well in 1986 and subsequently evaluated in the ravn-2 well in 1987. in 2010, the ravn-3 well was drilled to test the location of the oil–water contact and to evaluate the reservoir quality of the south-western flank of the field. oil was found at several intervals and the oil–water contact was located at a depth of 4572 m. the ravn member was deposited during an overall transgression of the heno plateau during the kimmeridgian. the member consists of up to 100 m thick marine shoreface deposits (johannessen 2010) where foreshore, middle and lower shoreface sandstones constitute the primary reservoirs (fig. 2). the sediments are strongly bioturbated and are dominated by very fineto fine-grained or muddy sandstones with occasional white, grey and light brown siltstones. methods sedimentological description of the ravn-3 core was made and 18 thin sections were prepared from samples from middle, lower and foreshore sandstones (fig. 2). petrographical investigations of the thin sections were undertaken with transmitted light microscopy. mineral abundances were quantified by point counting of minimum 500 grains. additional information was obtained from scanning electron microscopy (sem) of gold-coated rock chips and carbon-coated thin sections using a phillips xl 40 sem with a tungsten filament operating at 17 kv and 50–60 µa. porosity and permeability were measured on core plugs according to the api rp-40 standard (american petroleum institute 1998) at the geological survey of denmark and greenland. results the porosity and permeability of sandstone reservoirs reflect, among other things, depositional environmental, mineralogical composition and post-depositional diagenetic changes. in order to understand what affected porosity and permeability, these factors were investigated. detrital components – quartz is the dominant component in all sandstones. the feldspar group consists of k-feldspar and minor albite. k-feldspar is typically partially dissolved and diagenetic impact on reservoir sandstones of the heno formation in the ravn-3 well, danish central graben simone pedersen, rikke weibel, peter n. johannessen and niels h. schovsbo ringkøbing–fyn high feda graben ål basin 25 km salt structure normal fault reverse fault well r-1r-3 r-2 a 56°n 4°e salt dome province gert ridge nl g uk central graben mid north sea high mandal high outer rough basin inge high heno plateau tail end graben national border n structural high dk b fig. 1. a: present structural framework of the danish sector of the central graben. r-1, r-2, r-3: ravn-1, -2 and -3 wells. b: overview of the north sea area. green: land. modified from johannessen (2010). © 2018 geus. geological survey of denmark and greenland bulletin 41, 9–12. open access: www.geus.dk/bulletin http://www.geus.dk/bulletin 1010 minor mica, rock fragments and chlorite grains are present. accessory minerals are tourmaline, zircon and fe-ti oxides. detrital clay occurs as tangential coatings on detrital grains and as deformed clay clasts. diagenetic phases – sandstones are occasionally dominated by abundant sparry fe-dolomite and ankerite cement (fig. 3a; pedersen 2017). sporadic calcite inclusions occur enclosed in the fe-dolomite-ankerite cement. calcite from shell fragments was recognised in one sample. small amounts of fe-dolomiteankerite rhombs are present in samples where cement is not abundant. microcrystalline quartz coatings are common in several samples independent of depositional environment (fig. 3b). occasionally, excessive microcrystalline quartz also occurs in the intergranular pore space (fig. 3c). in a few sandstones, the detrital grain surfaces of quartz are only partly covered by microcrystalline quartz giving rise to growth of larger quartz overgrowths (fig. 3c). the amount of quartz overgrowths varies from 0.2 to 10.8 vol%. illite is present in all samples and depositional environments and occurs as fibrous and honeycomb-structured coatings (fig. 3d). authigenic illite occurs as protruding fibres growing from honeycomb-structured illiticsmectitic clay. illite fibres alternate with quartz overgrowths, and are at times enclosed in quartz overgrowth (fig. 3d). porosity versus permeability – the sandstones with highest porosity and permeability are dominated by microcrystalline quartz coatings and only little diagenetic illite is present together with a small amount of detrital clay (fig. 4; pedersen 2017). these sandstones are from the upper, middle and lower shoreface. two groups of sandstones are defined based on intermediate porosity and low to intermediate permeability. of these two groups, sandstones with quartz overgrowths and minor illite have slightly higher permeability than sandstones with microcrystalline quartz coatings and high illite and high detrital clay contents (fig. 4). these latter samples are from lower and middle shoreface. also the fe-carbonatecemented sandstones, which have the lowest porosity and permeability in the heno formation (fig. 4), represent lower and middle shoreface samples. comparison between the ravn-1, ravn-2 and ravn-3 wells – the ravn-3 well was correlated with the ravn-1 and ravn-2 wells based on available core and well log data (fig. 2). the various diagenetic phases in the ravn-3 well can be recognised in the other ravn wells. variations occur, such as quartz overgrowth and illitisation of detrital clay being more common in the ravn-1 well, compared to authigenic illite in the ravn-3 well, but the reservoir units can still be recognised. the variations seen in the ravn-1 cores are also present in the ravn-2 cores together with additional fractures filled with barite and ankerite. the porosity and permeability in the ravn-1 and ravn-2 wells lie within the same range as the sandstones in the ravn-3 well (fig. 4). cl si sand pbl lithology depositional environment sandstone clay or siltstone conglomerate lower shoreface middle shoreface foreshore structures disconnected wave ripples structureless due to bioturbation carbonate cemented sandstone thin section samples4619 m cl si sand pbl cl si sand pbl legend ravn-1 ravn-3 ravn-2 4280.77 m 4155 m 0 10 20 0.001 0.1 10 0 10 20 0 10 20 0.01 1 10 0.01 1 10 phi (%) kh (md) phi (%) kh (md) phi (%) kh (md) 0 20 40 60 80 100 120 h ei gh t a bo ve z er o (m ) offshore low angle cross stratification fig. 2. correlation panel of the cored parts of the ravn-1, -2, and -3 wells. the ravn-1 and ravn-2 logs are modified from johannessen (2010), whereas the ravn-3 core was logged for this study. the depositional environment described in the ravn-3 well (panterra 2011) is based on ichnofacies. phi: heporosity. kh: horizontal permeability. cl: clay. si: silt. pbl: pebble. 11 discussion early carbonate cement – intergrown sparry fe-dolomite and ankerite cement (fig. 3a) is interpreted to be sourced from dissolved calcite from shell fragments. calcite inclusions still occur between fe-dolomite and ankerite. this is supported by quartz grains appearing to be ‘floating’ in the carbonate cement, which indicates the previous presence of an early carbonate cement or fossils. fe-carbonates are considered more stable than calcite during late diagenesis and often replace earlier phases of carbonates (worden & burley 2003). early microcrystalline quartz – when early diagenetic microcrystalline quartz is present in the sandstones only minor quartz overgrowth has precipitated (fig. 3b). a biogenic opal ct phase, which has been dissolved without trace, may have resulted in supersaturated pore waters that sustained nucleation of microcrystalline quartz. grain-coating microcrystalline quartz has previously been proposed to preserve reservoir quality by impeding quartz overgrowth, which otherwise may occlude intergranular porosity and reduce permeability (aase et al. 1996; jahren & ramm 2000; weibel et al. 2010). the random growth of microcrystalline quartz may retard further development of both new microcrystalline quartz and quartz overgrowth (jahren & ramm 2000; weibel et al. 2010). when microcrystalline quartz does not fully cover detrital quartz, it cannot inhibit precipitation of quartz overgrowth (aase et al. 1996; weibel et al. 2010). quartz overgrowths – late diagenetic quartz overgrowths formed where the quartz grains were only partly covered by microcrystalline quartz. the quartz overgrowths probably formed under low silica oversaturation, which favoured less nucleation and promoted the growth of larger crystals (fig. 3c; jahren & ramm 2000). more intensive quartz ceca fe-do + an mq il mq qo mq + il mq il il mq 5 µm 20 µm 20 µm 50 µm a b c d fig. 3. a: abundant fe-dolomite (fe-do) and ankerite (an) occluding porosity and permeability. remnants of the original early calcite (ca) cement are present. b: random and abundant microcrystalline quartz (mq) coating detrital quartz grain, preventing quartz overgrowth (qo). note the fibrous illite (il). c: microcrystalline quartz on detrital quartz and in pore space together with authigenic illite. quartz overgrowth is partly enclosing microcrystalline quartz indicating that the quartz overgrowth precipitated later. d: abundant fibrous illite growing from honeycombstructured illite succeeding microcrystalline quartz and alternating with quartz overgrowth (qo). 10 1 0.1 0.01 0.001 0 5 10 15 2520 he porosity (%) g as p er m ea bi lit y (m d ) carbonate cemented microcrystalline quartz + illite + low detrital clay content microcrystalline quartz + illite + high detrital clay content thin section ravn-1 thin section ravn-2 thin section ravn-3 lower shoreface middle shoreface foreshore quartz overgrowth fig. 4. he porosity versus air permeability for all thin section samples from the ravn-3 well, together with data from the ravn-1 and ravn-2 wells. the thin section samples follow the trends from the ravn-3 well marked by the four ellipses, which depict the four characteristics of the diagenesis. the purple ellipse comprises samples dominated by microcrystalline quartz, illite and low detrital clay content. the green ellipse includes samples dominated by microcrystalline quartz, illite and high detrital clay content. the orange ellipse comprises samples dominated by quartz overgrowth and the blue ellipse by extensive sparry carbonate cement. 1212 mentation would have been expected in these quartz-rich sandstones (bjørlykke et al. 1989) as they have been buried to a depth of > 4 km and hence exposed to temperatures of 112–117°c as documented by vitrinite reflectance. as no stylolites were observed and as quartz overgrowth precipitated before and alternating with illite growth, another source for silica must have been present prior to transformation of smectite to illite. the continued precipitation of quartz overgrowth was probably from a silica source from the transformation of smectite to illite and dissolution of kfeldspar (hower et al. 1976; boles & franks 1979). this is supported by the honeycomb-structured smectite-illite coatings and partially dissolved detrital k-feldspar. illite – illite occurring as honeycomb structured coatings (fig. 3d) is a strong indicator of a smectite precursor (e.g. pollastro 1985). during burial, the percentage of illite in mixed-layer illite/smectite compared to smectite increases since smectite becomes more unstable with increasing temperature and pressure (pollastro 1985), which may be the reason why only illite is present in the ravn-3 well. the honeycomb-structured illite commonly forms nucleation or growth points for fibrous illite. k-feldspar is typically dissolved concomitantly with smectite dissolution, and k-feldspar can be an additional source for k+ and al3+ for further illite precipitation (hower et al. 1976; boles & franks 1979). the additional k+ and al3+ from the dissolution of k-feldspar might have led to further precipitation of the fibrous illite on illite honeycomb structures and singular precipitation in pore space. fe-dolomite-ankerite rhombs are considered a by-product of the transition from smectite to illite, which may liberate ca2+ and fe2+. conclusions the porosity and permeability of the reservoir sandstones in the ravn-3 well are controlled by the diagenetic phases formed during early and late diagenesis. the reservoir sandstones with the highest porosity and permeability are dominated by low to moderate amounts of microcrystalline quartz, illite and detrital clay. however, the more distal lower shoreface sandstones with the same dominating diagenetic phases, but with higher detrital clay content, are considered a poor reservoir due to low porosity and permeability. sandstones with dominance of quartz overgrowth and low detrital clay content have moderate to high porosity and low permeability. carbonate-cemented sandstones are considered non-reservoir due to insignificant porosity and low permeability. references aase, n.e., bjørkum, p.a. & nadeau, p.h. 1996: the effect of grain-coating microquartz on preservation of reservoir porosity. aapg bulletin 80, 1654–1673. american petroleum institute 1998: api recommended practice 40. recommended practices for core analysis, 240 pp. second edition. washington dc: api publishing services. bjørlykke, k., ramm, m. & saigal, g.c. 1989: sandstone diagenesis and porosity modification during basin evolution. geologische rundschau 78, 243–268. boles, j.r. & franks, s.g. 1979: clay diagenesis in wilcox sandstones of southwest texas: implications of smectite diagenesis on sandstone cementation. journal of sedimentary research 49, 55–70. hower, j., eslinger, e.v., hower, m.e. & perry, e.a. 1976: mechanism of burial metamorphism of argillaceous sediment: 1. mineralogical and chemical evidence. geological society of america bulletin 87, 725–737. jahren, j. & ramm, m. 2000: the porosity-preserving effects of microcrystalline quartz coatings in arenitic sandstones: examples from the norwegian continental shelf. in: worden, r.h. & morad, s. (eds): quartz cementation in sandstones. international association of sedimentologists special publication 29, 271–280. johannessen, p.n., dybkjær, k., andersen, c., kristensen, l., hovikoski, j. & vosgerau, h. 2010: upper jurassic reservoir sandstones in the danish central graben: new insights on distribution and depositional environments, 12–34. in: vining, b.a. (ed.): petroleum geology: from mature basins to new frontiers. proceedings of the 7th petroleum geology conference. geological society, london. panterra geoconsultants 2011: sedimentology, petrography and reservoir quality of cores from the ravn-3 well, north sea, denmark. geus archive report file no 28698. pedersen, s.s. 2017: the diagenetic impact on reservoir sandstones of the heno formation in the ravn-3 well, danish central graben, denmark. unpublished master thesis, university of copenhagen. pollastro, r.m. 1985: mineralogical and morphological evidence for the formation of illite at the expense of illite/smectite. clays and clay minerals 33, 265–274. taylor, t.r., giles, m.r., hathon, l.a., diggs, t.n., braunsdorf, n.r., birbiglia, g.v., kittridge, m.g., macaulay, c.i. & espejo, i.s. 2010: sandstone diagenesis and reservoir quality prediction: models, myths, and reality. aapg bulletin 94, 1093–1132. weibel, r., friis, h., kazerouni, a.m., svendsen, j.b., stokkendal, j. & poulsen, m.l.k. 2010: development of early diagenetic silica and quartz morphologies – examples from the siri canyon, danish north sea. sedimentary geology 228, 151–170. worden, r. & burley, s. 2003: sandstone diagenesis: the evolution of sand to stone. in: burley, s.d. & worde, r.h. (eds): sandstone diagenesis: recent and ancient. international association of sedimentologists special publication 4, 3–44. authors’ addresses s.p., university of copenhagen, department of geosciences and natural resource management, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: simonepeder89@gmail.com. r.w., n.h.s. & p.j., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. mailto:simonepeder89@gmail.com the jurassic of denmark and greenland 265 the jurassic deposits in the danish central graben are of particular interest in relation to hydrocarbon exploration, as they include units of reservoir sandstones as well as oiland gas-producing source rocks (fig. 1). hydrocarbon discoveries have been made in both middle and upper jurassic sandstones, and the middle jurassic harald field is under production. the distribution of both reservoirs and source rocks reflects the complex sequence stratigraphy of the jurassic of the danish central graben jan andsbjerg and karen dybkjær a sequence stratigraphic framework is established for the jurassic of the danish central graben based primarily on petrophysical log data, core sedimentology and biostratigraphic data from about 50 wells. regional seismic lines are used to assist in the correlation of some wells and in the construction of isochore maps. in the lower jurassic (hettangian–pliensbachian) succession, five sequences have been identified. the middle jurassic is subdivided into four sequences that together span the uppermost aalenian/lowermost bajocian to the callovian. in the upper jurassic, better well coverage permits greater stratigraphic resolution, and 11 sequences are identified and mapped. on the basis of the sequence stratigraphic correlation and the construction of isochore maps for individual sequences, the jurassic basin history of the danish central graben can be subdivided into seven discrete phases: (1) shallow marine and offshore sediments deposited in a prerift basin extending from the north sea to the fennoscandian border zone (hettangian–pliensbachian). (2) uplift and erosion in association with a toarcian–aalenian north sea doming event. a major hiatus represents this phase in the study area. (3) terrestrial and marginal marine sedimentation during initial rifting (latest aalenian/earliest bajocian – late callovian). (4) early oxfordian – early kimmeridgian transgression during and after a rift pulse. the sedimentary environment changed from coastal plain and marginal marine to fully marine. (5) regression associated with a cessation or slowing of subsidence during a structural rearrangement that took place in the late kimmeridgian during a break in the main rift climax. shallow to marginal marine sandstones were deposited above an erosion surface of regional extent. (6) deep-water mudstones deposited in a composite graben with high subsidence rates related to rift pulses (latest late kimmeridgian – middle middle volgian). (7) deposition of organic-rich mudstones and turbidite sandstones during the late middle volgian – early ryazanian. the main basin shallowed, became more symmetrical and experienced a decreasing rate of subsidence, recording the onset of the post-rift stage. a relative sea-level curve is constructed for the middle–late jurassic. it shows close similarity to published eustatic (global) and relative (north atlantic area) sea-level curves in the latest bathonian – late early kimmeridgian, but differs in the late kimmeridgian – middle volgian interval, probably due to the high rate of subsidence in the study area. keywords: danish central graben, jurassic, sequence stratigraphy, palaeogeography, basin evolution, sea-level changes geological survey of denmark and greenland, geocenter copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ja@geus.dk geological survey of denmark and greenland bulletin 1, 265–300 (2003) © geus, 2003 tectonic evolution of the area and a depositional history strongly influenced by relative sea-level changes. data from a large number of wells in the study area have been released in recent years, contributing important new information to this study. the aim of this paper is to present a sequence stratigraphic model for the jurassic deposits of the danish central graben. 266 ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■ ■ ▲ ▲ ▲ ■■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■■■■ ■■ ■■ ■■ ■ ■ ■■ ■ ■ ■ ■ ■■ ■■ ■ ■ ■ ■ ■■ ■■ ■ ■ ■■ ■■ ■■ ■ ■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■■ ■■ ■ ■ emma-1 cleo-1 22 1 karl-1 tordenskjold-1 gulnare-1 i-1 elin-1 deep adda-1 v-1 ugle-1 m-8 u-1 deep gorm-1 anna-3 o-1 2/12-12/11-7 baron-2 amalie-1 ravn-2 ravn-1 w-1 eg-1 liva-1 b-1 falk-1 elly-1 elly-2 elly-3 edna-1 jens-1 n. jens-1 bo-1 e-1 adda-1 alma-1 se. igor-1 g-1 kim-1 lone-1 diamant-1 gwen-2 q-1 jeppe-1 mona-1 nora-1 iris-1 3/7-4 lulu-1 2/12-2 gert-31 4 w. lulu-3 john flank-1 s øgne b asin piggvar terrace gert ridge feda graben inge high heno plateau m ads high tail end g raben a rne–elin g raben poul plateau r o sa b asin gertrud plateau/graben mandal high ål basin o uter rough basin c offee soil fault salt dome province mid north sea high ringkøbing–fyn high (east north sea block) reverse fault triangles showing dip of fault plane normal fault boxes showing dip of fault plane salt structure well ▲ ■■ 20 km n 4°e 5°e 24 56°30' 56°n 500 km uk central graben moray firth basin viking graben norway denmark germany the netherlands fig. 1. maps showing the location of the study area (dark blue) within the jurassic north sea rift system, the jurassic structural elements in the danish central graben and the position of the wells incorporated in the study; note that all wells forming the database for the isochore and palaeogeographic maps are indicated, although some are not specifically mentioned in the text or illustrated. the middle jurassic harald field is delimited by the location of the west lulu-1, -2 and -3 wells; note that the structural feature denoted the gertrud plateau/graben is so indicated to reflect the structural evolution of this region in the late jurassic, from a positive plateau in the early late jurassic to an actively subsiding graben from the early volgian. regional geology the central graben forms the southern part of the north sea rift system. active rifting took place in this area both in the triassic and in the late middle jurassic – late jurassic (roberts et al. 1990; ziegler 1990). the 500 km long central graben, approximately 150 km of which are situated in the danish north sea sector, consists of a nnw–sse-trending complex of half-grabens, and a subordinate n–s-trending segment to the south extending into the german and dutch sectors. the main bounding fault in the danish sector is the coffee soil fault, which forms the eastern margin of the danish central graben (fig. 1). the early jurassic was characterised by a slow relative sea-level rise and deposition of marine muds of the fjerritslev formation, probably over most of the danish area (fig. 2; larsen 1966; michelsen 1978, 1989; pedersen 1986; michelsen et al. 1987). in latest early jurassic – earliest middle jurassic times, the whole area was uplifted and most of the lower jurassic section was removed by erosion (hallam & sellwood 1976; ziegler 1982, 1990; underhill & partington 1993). the uplift may have occurred in response to the development of a rift dome extending 700 km in a north–south direction and 1000 km east–west across the central north sea (whiteman et al. 1975; eynon 1981; ziegler 1990; underhill & partington 1993). domal uplift appears to have been accompanied by the development of a volcanic complex at the triple junction between the viking graben, the central graben and the moray firth basin (ziegler 1990). sedimentation resumed in the danish part of the central graben during the middle jurassic, with deposition of the sandstone-dominated bryne and lulu formations and the mudstonedominated middle graben formation (michelsen et al. 2003, this volume). these sediments are restricted to the søgne basin, the tail end graben and the salt dome province (vollset & doré 1984; jensen et al. 1986). the first marine transgression in the danish central graben occurred during the callovian–oxfordian, probably reflecting the onset of domal collapse combined with eustatic sea-level rise (ziegler 1990; underhill & partington 1993). 267 u pp er volgian kimmeridgian oxfordian callovian bathonian bajocian aalenian toarcian pliensbachian sinemurian hettangian sand-dominated formations m id dl e lo w er s n nw se poul fm lola fm fjerritslev fm farsund fm ju ra ss ic middle graben fm marine mud-dominated formations paralic mud-dominated formationorganic-rich mudstones missing section farsund fm heno fm lulu fm bryne fm heno fm bo mbbo mb ryazanian chronostratigraphy l. cret. fjerritslev fm lola fm middle graben fm bryne fm fig. 2. jurassic lithostratigraphy of the danish central graben, based on jensen et al. (1986) and michelsen et al. (2003, this volume). l. cret., lower cretaceous. during late jurassic time, the feda graben, the heno plateau and the gertrud graben became actively subsiding depositional basins, and the depositional area was later extended to the outer rough and ål basins (fig. 1). deep-water conditions were established in the danish central graben during the oxfordian–kimmeridgian, when the marine mudstones of the lola formation were deposited. during the late kimmeridgian, shallow marine sandstones of the heno formation were deposited on plateau areas. the transgression culminated during the volgian with deposition of the deep marine mudstones of the farsund formation (vollset & doré 1984; jensen et al. 1986). the early cretaceous saw a change to more passive thermal subsidence (roberts et al. 1990). in addition to rift tectonics, basin development in the central graben was also strongly influenced by the presence of mobile zechstein salt. salt movements had a profound influence on the development of depocentres in the søgne basin, the tail end graben and the salt dome province (korstgaard et al. 1993; sundsbø & megson 1993). stratigraphic methods the stratigraphic analysis is based on data from about 50 wells that penetrate jurassic deposits in the danish central graben, and a few wells from the norwegian north sea sector located near the boundary (fig. 1). well logs, including gamma-ray, sonic, neutron/density and resistivity logs, sedimentological core logs, lithology logs, and biostratigraphic data were used in the study. furthermore, a number of interpreted regional seismic lines were included to support the generation of isochore maps. biostratigraphy the biostratigraphic data used in this study are confined to palynomorphs because information from other groups of microfossils is too scarce in most wells to be useful for detailed correlation. the biostratigraphic correlations are based on events rather than recognition of biozones. the events used in our study are mainly last occurrence datums (i.e. first downhole appearance) of dinoflagellate cyst species (fig. 3). the biostratigraphic information utilised in the study includes published data (birkelund et al. 1983; hoelstad 1986; poulsen 1986, 1991; heilmann-clausen 1987; johannessen et al. 1996), unpublished reports of the geological survey of denmark and greenland and of industrial service companies, and the results of new analyses made specifically for this study. due to poor core coverage, being generally restricted to sandy successions of relatively few wells, most data were derived from cuttings samples. although sample quality in the lower jurassic succession was variable, it was possible to establish a palynostratigraphic framework based on a small number of events and stratigraphically diagnostic palyno-assemblages (fig. 3a). a detailed palynostratigraphical subdivision of the middle jurassic succession is precluded by the sparse occurrence of age-diagnostic palynomorphs (fig. 3b); although scanty, the available records made it possible to give an approximate age for the sequences and major events in the middle jurassic. in the upper jurassic, the large sediment thicknesses, the somewhat better data quality, the favourable environment, and the higher diversity of dinoflagellate cyst species probably all contribute to a higher biostratigraphic resolution (fig. 3c). sequence stratigraphy the sequence stratigraphic terminology applied is that introduced by the exxon group (posamentier et al. 1988; posamentier & vail 1988; van wagoner et al. 1988, 1990). sequences are subdivided into lowstand, transgressive and highstand systems tracts. a four-fold subdivision with the addition of the forced regressive systems tract (hunt & tucker 1992, 1995) is not attempted due to the scale of the study and the nature of the data. construction of the sequence stratigraphic framework is based on correlation of well log patterns, combined with biostratigraphic datings. the most conspicuous log patterns, supported by the most important biostratigraphic events, were used to produce a coarse grid, within which more detailed correlations were made. most commonly the gamma-ray logs were used for high-resolution correlations but resistivity logs formed the primary tool in some organic-rich mudstone units. distinct sonic log markers were occasionally used for correlation. maximum flooding surfaces (mfs) are considered the most reliable correlation surfaces in successions dominated by marine mudstones, whereas sequence boundaries (sb), and occasionally flooding surfaces (fs), are of primary importance in silty and sandy units. both sb and mfs key surfaces can normally be traced through most of the basin. in sections dominated by marine mudstones and siltstones, the sb is usually picked at the turnaround point between a coars268 269 age dapcodinium priscum cerebropollenites macroverrocosus2) liasidium variabile liasidium variabile dapocodinium priscum5) boreal standard zones bioevents (palynology) first appearance datums last occurrence datums time in ma ep oc h ea rl y ju ra ss ic 190 195 185 200 205 la te t ri as si c (gradstein et al. 1994) la te la te l la te la te ea rl y ea rl y m id dl e ea rl y ea rl y e m id dl e dispansum pseudoradiosa aalensis thouarsense variabilis bifrons falciferum tenuicostatum spinatum margaritatus davoei ibex jamesoni raricostatum oxynotum obtusum turneri semicostatum bucklandi angulata liasicus planorbis marshi reticulatus rhaetian hettangian sinemurian pliensbachian toarcian first appearance last occurrence tie to boreal standard zones uncertain a fig. 3. bioevents used for dating the sequences in the lower jurassic (a), middle jurassic (b) and upper jurassic – lowermost cretaceous (c) of the danish central graben. the tie to boreal standard zones are according to riding & thomas (1992), unless otherwise indicated: 1, fenton & riding (1987); 2, dybkjær (1991); 3, poulsen (1991); 4, costa & davey (1992); 5, poulsen (1992); 6, poulsen & riding (2003, this volume). note in fig. 3c: (1) the species indicated with an asterisk (muderongia simplex) was previously referred to muderongia sp. a davey 1979; see further discussion in poulsen (1996). (2) although the lod is indicated at the base of the kerberus zone, the species senoniasphaera jurassica (double asterisk) may occur sporadically up to the anguiformis zone. e, early; l, late. ening-upwards unit and a fining-upwards unit, or at the sharp base of a relatively coarse-grained bed occupying the position of the turnaround point. well logs form the primary tool for the identification of lithologies, sedimentary facies and sedimentary successions, and thus for the identification of key surfaces and systems tracts. the cored sections have been the subject of more detailed sedimentological facies analysis and palynofacies investigations (hoelstad 1986; johannessen & andsbjerg 1993; johannessen et al. 1996; andsbjerg 2003, this volume; ineson et al. 2003, this volume; johannessen 2003, this volume); these results have been used to support the sequence stratigraphic subdivision. the lower jurassic succession shows laterally consistent log patterns and, where a unit has not been removed by erosion, correlation of key surfaces is possible over wide areas. in the middle jurassic section, only two or three key surfaces can be correlated from one subbasin to another. a more detailed, four-fold sequence subdivision of the middle jurassic deposits within the søgne basin is presented by andsbjerg (2003, this volume). the laterally uniform log patterns of most of the upper jurassic succession allow detailed well correlations. in addition to the gamma-ray log, the resistivity logs proved to be a robust tool for high-resolution correlations in the lower volgian succession of the 270 lamberti athleta coronatum jason calloviense herveyi koenigi discus orbis hodsoni morrisi subcontractus progracilis tenuiplicatus zigzag parkinsoni garantiana subfurcatum humphriesianum sauzei laeviuscula discites concavum murchisonae opalinum callovian bathonian bajocian aalenian impletosphaeridium varispinosum adnatosphaeridium caulleryi impletosphaeridium varispinosum kekryphalospora distincta1) 180 170 165 175 160 la te la te la te l ea rl y ea rl y ea rl y e m id . m id dl e m id dl e age boreal standard zones bioevents (palynology) first appearance datums last occurrence datums time in ma (gradstein et al. 1994) m id dl e ju ra ss ic ep oc h first appearance last occurrence tie to boreal standard zones uncertain b southern feda graben and the northernmost heno plateau. most key surfaces in the upper jurassic succession are picked in marine mudstones and siltstones. of these key surfaces, only one sequence boundary (the base kimm-2 sb) shows any sign of significant erosion. unusually high organic contents in the upper volgian – ryazanian mudstones of the bo member (michelsen et al. 2003, this volume) probably influence log responses and impede sequence stratigraphic interpretation (ineson et al. 2003, this volume). in the uppermost sequence of this study, ryaz-1, a maximum flooding surface is not identified. it might have been picked at the highest gamma-ray peak in accordance with conventional sequence stratigraphic concepts, but in this case all of 271 age boreal standard zones bioevents (palynology) first appearance datums last occurrence datums time in ma ep oc h (gradstein et al. 1994) 150 155 140 145 albidum stenomphalus icenii kochi runctoni lamplughi preplicomphalus primitivus oppressusanguiformis kerberus okusensis glaucolithus albani fittoni rotunda pallasioides pectinatus hudlestoni wheatleyensis scitulus elegans autissiodorensis eudoxus mutabilis cymodoce baylei rosenkrantzi regulare serratum glosense tenuiserratum densiplicatum cordatum mariae la te ju ra ss ic ea rl y c re ta ce ou s ea rl y ea rl y ea rl y ea rl y la te la te la te la te m id dl e m id dl e dingodinium spinosum4) dichadogonyaulax culmula4) rotosphaeropsis thula4) egmontodinium expiratum4) egmontodinium polyplacophorum dichadogonyaulax pannea glossodinium dimorphum muderongia simplex* gochteodinia mutabilis senoniasphaera jurassica6)**scriniodinium inritibile occisucysta balia perisseiasphaeridium pannosum oligosphaeridium patulum cribroperidinium longicorne subtilisphaera? paeminosa endoscrinium luridum stephanelytron scarburghense endoscrinium galeritum scriniodinium crystallinum nannoceratopsis pellusida compositosphaeridium polonicum rigaudella aemula wanaea spp. ctenidodinium continuum subtilisphaera? paeminosa3) s.? inaffecta3) oxfordian ryazanian volgian kimmeridgian first appearance last occurrence tie to boreal standard zones uncertain paratolliavalanginian c the bo member may form part of the transgressive systems tract as suggested for this type of deposits by posamentier & james (1993). this level marks the top of the studied succession. in those wells where the uppermost part of the farsund formation is preserved, it is dated as late ryazanian. sequence stratigraphic subdivision and basin history the jurassic succession in the danish central graben is subdivided into 20 sequences, referred to according to their age (i.e. hettangian sequences: hett-1, hett-2). the main characteristics of each sequence – log expression of sequence boundaries, lithology, thickness, age – are summarised and illustrated below. the lower boundary is described for each sequence, the upper boundary being described under the succeeding sequence. lithologies in the sequence descriptions have been interpreted primarily on the basis of gamma-ray logs, supported by other log types and by core data when available. in the sandstone-dominated intervals, increasing gamma-ray values are interpreted to reflect finingupwards grain-size trends, and decreasing values to reflect coarsening-upwards trends. in the mudstone successions, changing gamma-ray values are interpreted to reflect changes in clay content; increasing gammaray values are thus interpreted to record increasing clay content and decreasing gamma-ray values to record decreasing clay content. thin carbonate beds are represented by low-value gamma-ray spikes. mudstone intervals with exceptionally high gamma-ray values are interpreted as being particularly rich in organic matter. based on these inferred relationships between gammaray values and lithologies, increasing-upwards gammaray patterns are generally interpreted to be characteristic of the transgressive systems tract (tst) and decreasingupwards gamma-ray patterns of the highstand systems tract (hst). the maximum flooding surface (mfs) is picked at the turnaround point where gamma-ray values are at a maximum between the increasing gammaray values of the tst below and decreasing gamma-ray values of the hst unit above. the identified sequences are of approximately the same order of thickness. due to variations in subsidence rate, the duration of the sequences recognised in the lower and middle jurassic is somewhat longer than for the upper jurassic sequences. all sequences, however, probably correspond to 3rd order sequences as defined by vail et al. (1977) and van wagoner et al. (1990). the evolution of the danish central graben can be considered in terms of seven depositional phases on the basis of the subdivision of the jurassic deposits into sequences and systems tracts. the significant influence of tectonics on depositional patterns in rift settings has been demonstrated by a number of studies (surlyk & clemmensen 1983; rosendahl 1987; gabrielsen et al. 1990; prosser 1993; nøttvedt et al. 1995). however, we have attempted to relate the development of the depositional phases to the effects of both tectonics and eustasy. the nomenclature and subdivision of rifting phases introduced by prosser (1993) and adapted by nøttvedt et al. (1995) are used in an attempt to relate the depositional phases of the present study to rifting stages. the proto-rift stage is characterised by flexural subsidence and interrupted by domal uplift. the subsequent rift stage is subdivided into rift initiation and rift climax: rift initiation is associated with fault block compartmentalisation and weak tilting whereas the rift climax is characterised by the maximum rate of fault displacement and strong tilting. the early post-rift stage is marked by the cessation of significant activity at the main faults and a declining rate of regional subsidence. it should be appreciated, however, that the three stage model of nøttvedt et al. (1995) refers to the evolution of rift basins in general. in the present study, the model has been applied to minor basins in a large rift structure and the evolution of these basins may not be representative of the rift structure as a whole. this is emphasised by the diachronous onset of subsidence in neighbouring basins, as described below. in the present study, several pulses of active faulting are inferred to have taken place between the initiation of the middle jurassic rift stage and the early cretaceous early post-rift stage. we have followed surlyk (1978, 1989) and blair & bilodeau (1988) in correlating overall fining-upwards successions in rift basins with periods of active tectonic subsidence, and large-scale coarsening-upwards successions with periods of relative tectonic quiescence. phase 1. pre-rift shallow marine deposition (hettangian–pliensbachian) sequence definitions hett-1 sequence the hett-1 sequence is present only in the deep gorm-1, m-8 and u-1 wells; it consists of marine mudstones. the lower sequence 272 boundary is located at the change from the low gamma-ray values of the underlying triassic deposits to the high gamma-ray values, characteristic of the mudstones of the hett-1 sequence (figs 4, 5). the boundary may be abrupt or gradational. the thickness of the sequence attains a maximum of 23 m in m-8. stratigraphically useful bioevents have not been recorded from this unit. the sequence is thought to be of hettangian age due to its stratigraphic position in the lowermost part of the fjerritslev formation, unconformably overlying upper triassic deposits. hett-2 sequence the marine mudstones of the hett-2 sequence have only been found in six wells but are considered to extend throughout the salt dome province. the boundary between the hett-1 and hett-2 sequences is located at a shift to mudstones characterised by somewhat lower clay contents (figs 4, 5). in wells where the hett-2 sequence directly overlies pre-jurassic sediments (e.g. edna-1; fig. 5), the lower sequence boundary is marked by an abrupt lithological shift. the thickness of the sequence shows very little variation, attaining a maximum of 64 m in edna-1. the abundance of bisaccate pollen of pinuspollenites minimus within this sequence in the edna-1, john flank-1 and o-1 wells indicates the presence of the pinuspollenites–trachysporites zone of lund (1977). this zone is known from the danish basin, skåne and northern germany and is referred to the hettangian (lund 1977; dybkjær 1991). a hettangian age is further supported by the occurrence of the dinocyst species dapcodinium priscum in the wells mentioned above and the absence of the pollen species cerebropollenites macroverrucosus (fig. 3a). sin-1 sequence the marine mudstones of this sequence probably extend throughout the salt dome province. the lower sequence boundary is typically located at a marked shift to mudstones with higher clay contents (figs 4, 5). the thickness of the sequence attains a maximum of 63 m in deep gorm-1. the last occurrence datum (lod) of dapcodinium priscum in the lower part of the sequence in the o-1 well, and the first appearance datum (fad) of cerebropollenites macroverrucosus in core samples from the middle part of the sequence in the deep gorm-1 well, indicate a latest hettangian – early sinemurian age for the lower part of the sequence. in combination with the lod of liasidium variabile close to the upper boundary of this sequence in o-1, this indicates a latest hettangian – sinemurian age for the sequence (fig. 3a). pliens-1 sequence this sequence is only present in the deep gorm-1 and edna-1 wells situated in the north-western part of the salt dome province. in both wells, it is represented by marine mudstones. the lower sequence boundary is located at the top of the well-defined coarsening-upwards hst of the sin-1 sequence (figs 4, 5). the thickness of the sequence attains a maximum of 63 m in edna-1. the common occurrence (in sidewall cores) of pollen referred to the genus chasmasporites indicates the presence of the chasmasporites zone (koppelhus & nielsen 1994), defined in the korsodde section, bornholm, denmark. the chasmasporites zone is referred to the pliensbachian, an age assignment supported by the stratigraphic position of this sequence immediately above the lod of liasidium variabile. the acme of the small, spherical pollen referred to the genus spheripollenites, characteristic of the uppermost pliensbachian – lower toarcian deposits in the danish basin and known from many other locations in northwest europe (dybkjær 1991), has not been reported from the danish central graben. this may indicate that deposits of latest pliensbachian – toarcian age are absent in the study area. pliens-2 sequence the marine mudstones of this sequence have only been recognised in the deep gorm-1 and edna-1 wells (fig. 4). the sequence boundary is positioned where the lithology changes abruptly from silty mudstones of the pliens-1 hst to more clay-rich mudstones (fig. 5). in the deep gorm-1 well, the sequence is represented by less than 10 m of clay-rich mudstone; in edna-1, the sequence attains a thickness of 62 m. the common occurrence (in sidewall cores) of chasmasporites sp. pollen in edna-1 and the absence of a spheripollenites sp. acme and of other toarcian bioevents (such as the fad of the spore genera leptolepidites, ischyosporites, manumia or staplinisporites), indicate a pliensbachian age for this sequence. basin evolution the early jurassic was a tectonically quiescent period between phases of active rifting in the triassic and in the middle–late jurassic (cartwright 1991; nøttvedt et al. 1995). a eustatic rise in sea level during the early jurassic was proposed by hallam (1981) and haq et al. (1988), compatible with observations from the danish basin (michelsen 1978, 1989; pedersen 1986; dybkjær 1988, 1991). uniform lithologies, dominated by shelf mudstones, indicate that marine conditions existed across a major north sea basin and extended into the danish basin (michelsen 1978, 1982, 1989; michelsen et al. 1987; pedersen 1985; nielsen 2003, this volume). the lower jurassic marine mudstones within the study area have only been found in the salt dome province. they were probably deposited over a wider area, but were removed by erosion during the regional uplift phase (phase 2) at the early–middle jurassic transition (andersen et al. 1982; gowers & sæbøe 1985; underhill & partington 1993). the extent of lower 273 274 k im m -1 o x2 * * m id dl e ju ra ss ic , u nd iff er en tia te d pl ie ns -2 pl ie ns -1 si n1 h et t2 h et t2 h et t1 c al -1 o x1 si n1 ba j-1 – ba t1 c al -1 – ?k im m -1 g r so ni c 41 64 m r es g r so ni c 39 35 m r es g r so ni c 34 30 m r es g r so ni c 23 14 m r es g r so ni c 31 55 m ed na -1 d ee p g or m -1 u -1 jo hn f la nk -1 o -1 r es 20 k m o -1 jo hn f la nk -1d ee p g or m -1 u -1 ed na -1 m ar in e m ud st on e m ar in e si lts to ne m ar in e an d pa ra lic sa nd st on e fl uv ia l s an ds to ne fl oo dp la in m ud st on e an d si lts to ne 10 0 mm ax im um fl oo di ng s ur fa ce se qu en ce b ou nd ar y jurassic deposits in the deepest parts of the tail end graben and the søgne basin is unknown due to the lack of wells penetrating the base of the middle jurassic in these areas. five transgressive–regressive cycles, corresponding to the hett-1 – pliens-2 sequences are identified in the lower jurassic succession (e.g. deep gorm-1, edna-1; fig. 4). individual sequences show a remarkable lateral consistency in thickness (fig. 4), indicating a uniform subsidence history in the study area during most of the early jurassic. phase 2. uplift and erosion in the proto-rift phase (latest pliensbachian – latest aalenian) basin evolution a major unconformity separates lower jurassic deposits from the overlying middle and upper jurassic deposits over much of the north sea area. it has been suggested that this unconformity was caused by domal uplift in the toarcian to aalenian, centred at the triple junction between the viking graben, the moray firth basin and the central graben (whiteman et al. 1975; hallam & sellwood 1976; eynon 1981; ziegler 1982; underhill & partington 1993). the unconformity may therefore be an example of a proto-rift unconformity, described by nøttvedt et al. (1995) as being typical of rift systems with thermally-induced domal uplift before or at the onset of active stretching. as a result of the regional uplift, most of the lower jurassic deposits were removed. in the southern part of the danish central graben, where lower jurassic deposits are preserved, differential erosion of the upper part of the succession may be due to local salt tectonics. regional erosion caused a larger part of the succession to be removed over the top of the salt structures. the full lower jurassic succession shows thickness variations from about 60 m in u-1 to about 260 m in edna-1 (fig. 4). reworked lower jurassic palynomorphs recorded from the upper kimmeridgian sandstones of the heno 275 facing page: fig. 4. sequence stratigraphic correlation of the lower, middle and lowermost upper jurassic in the salt dome province. the lower jurassic succession is truncated by an unconformity at the base of the middle jurassic. the depth given beneath each well in this and subsequent correlation panels indicates the depth of the lowermost tieline (typically base-jurassic section) below well reference level; for those few wells that did not penetrate the jurassic section, the depth indicates the base of the logged section (i.e. well total depth). the log suite utilised in the study and illustrated in this and subsequent figures includes gamma-ray (gr), sonic velocity (sonic) and resistivity (res) logs. 50 m gr sonic res pliens-2 pliens-1 sin-1 hett-2 3864 m b. msl. fj er ri ts le v fm d ud ge on sa lif er ou s fm lo w er ju ra ss ic t ri as si c h et ta ngi an si ne m ur . pl ie ns ba ch ia n m. jur. sequences lithostratigraphy chronostratigraphy lithostratigraphy chronostratigraphy edna-1 gr sonic res cal-1 pliens-1 sin-1 hett-2 3698 m b. msl. fj er ri ts le v fm lo w er ju ra ss ic h et ta ng ia n si ne m ur . pl ie ns ba ch ia n sequences deep gorm-1 hett-1 m . ju r. c al lo vi an m . g ra be n fm bryne fm pliens-2 maximum flooding surface sequence boundary fig. 5. lower jurassic sequences (hett-1 – pliens-1) in the deep gorm-1 and edna-1 wells. for reference purposes, the depth (in metres below mean sea level) is indicated for a selected sequence boundary, typically in the upper levels of the illustrated section (see also figs 6, 10, 12, 19, 22). jur., jurassic; m., middle; sinemur., sinemurian. formation in the northern part of the danish central graben (gwen-2; johannessen et al. 1996) supports the assumption that lower jurassic deposits were significantly more extensive prior to mid-jurassic erosion. phase 3. terrestrial and marginal marine deposition during the rift initiation stage (latest aalenian/earliest bajocian – late callovian) sequence definitions aalen-1 sequence this sequence is encountered in most wells drilled in the søgne basin and in the nora-1 well in the tail end graben (figs 6, 7). it lies unconformably on either pre-jurassic or lower jurassic deposits. the tst dominates the sequence and consists of a number of minor, backstepping, fining-upwards sandstone intervals, interpreted as stacked fluvial channels, with subordinate floodplain or lacustrine mudstones. the mfs is a distinct gamma-ray high, situated close to the strongly erosional upper sequence boundary. the thickness of the sequence attains a maximum of 69 m in west lulu-1 (fig. 6). no stratigraphically useful bioevents were recorded from this sequence, but the lack of toarcian bioevents (as discussed above with respect to the pliens-1 and pliens-2 sequences) and the lod of kekryphalospora distincta in the sequence above, may be seen as indirect evidence of an aalenian or earliest bajocian age. baj-1 sequence the baj-1 sequence extends throughout the søgne basin and the tail end graben (nora-1), and further south to the northern and eastern parts of the salt dome province (e.g. alma-1, o-1). it may also be present in structurally deep locations elsewhere in the salt dome province. the lower sequence boundary is a pronounced erosional surface (figs 4, 6, 7). the lower part of the sequence typically consists of two laterally extensive channel sandstones separated by a fining-upwards/coarsening-upwards fine-grained interval (e.g. amalie-1; fig. 6). the channel sandstone interval may be interpreted as a lowstand systems tract (lst). overlying this interval is a fining-upwards tst (amalie-1, west lulu-1; fig. 7). the hst consists of floodplain sandstones and mudstones. the thickness of the sequence attains a maximum of 99 m in amalie-1. the lod of kekryphalospora distincta in the upper part of this sequence in the alma-1 well indicates an aalenian or earliest bajocian age for the sequence (fig. 3b). bat-1 sequence this sequence is present in all wells that encountered middle jurassic deposits in the søgne basin, the tail end graben, the salt dome province and the southernmost heno plateau (figs 6, 7). in most wells, the lower sequence boundary is located at the base of a finingupwards channel sandstone (e.g. west lulu-1; fig. 6). the mfs is located in a mudstone interval, several metres thick. the hst is rarely present due to erosion; where present (e.g. west lulu-1; fig. 6), it consists of a coarsening-upwards interval of interbedded mudstones and sandstones. the thickness of the sequence attains a maximum of 64 m in west lulu-1. the occurrence (in a core sample) of adnatosphaeridium caulleryi in the lower part of this sequence in west lulu-1 indicates an age no older than bathonian (fig. 3b). in combination with the occurrence of impletosphaeridium varispinosum immediately above the upper boundary of the sequence in west lulu-3, this indicates a possible age range for this sequence of latest bajocian – earliest callovian. cal-1 sequence this sequence is present in areas where the bat-1 sequence is also found (fig. 8a). in most wells, the lower sequence boundary is a very distinct erosional surface (see amalie-1; fig. 6), which locally shows truncation of at least 10–20 m (e.g. west lulu-3; fig. 7). the sequence boundary is overlain by a section of fluvial to estuarine channel sandstones up to 30 m thick (e.g. amalie-1), which is interpreted as the fill of an incised valley. this interval probably represents the lst and the lowermost part of the tst. in the søgne basin and the tail end graben, the remainder of the tst consists of a succession of sandstone-dominated, paralic to shallow marine deposits, overlain by fully marine mudstones (fig. 7). the paralic deposits include a number of coal beds, up to 3 m thick. the hst consists of a rather indistinct coarsening-upwards interval of marine mudstones. the sequence has a depocentre in the northern part of the tail end graben (fig. 8a). in the nora-1 well, where the sequence includes a section that contains tuffaceous sediments (4986–4932 m), it attains a thickness of 288 m. in other wells, the sequence does not exceed 100 m in thickness. the lod of impletospaeridium varispinosum in the lower part of this sequence in west lulu-3 combined with the lod of ctenidodinium continuum in the sequence above (ox-1) in u-1 and west lulu-1 indicate an early callovian to earliest oxfordian age for this sequence. the occurrence of wanaea acollaris and w. thysanota in the middle part of the sequence in the john flank-1 and west lulu-1 wells supports this age assignment (fig. 3a). basin evolution deposition resumed in the danish central graben with the accumulation of uppermost aalenian or lowermost bajocian terrestrial sediments after the regional uplift and possible incipient collapse of the domal structure. subsequent middle jurassic deposits rest unconformably on pre-jurassic and lower jurassic sediments (fig. 4). 276 the initiation of syn-rift subsidence is demonstrated by an asymmetric distribution of the sediment package in the søgne basin as shown by seismic evidence (møller 1986, fig. 5; cartwright 1987; korstgaard et al. 1993, figs 39, 40) and well log data (andsbjerg 2003, this volume). salt tectonics also influenced sediment distribution and the location of local depocentres in the søgne basin, the tail end graben and the salt dome province (mogensen et al. 1992; korstgaard et al. 1993). middle jurassic deposits are preserved in the søgne basin, the tail end graben and the salt dome province with a depocentre located to the east near the coffee soil fault (figs 4, 7). sandstones and mudstones of sequences aalen-1 – bat-1 were deposited in a fluvially dominated environment during the early part of the middle jurassic (latest aalenian/earliest bajocian – late bathonian; fig. 9a). deeper parts of the basins were periodically inundated by lakes. lacustrine conditions in the central parts of the basins were coeval with wet floodplain conditions in marginal locations. the occurrence locally of rare marine palynomorphs and tidal indicators such as flaser bedding and double mud-drapes suggest that deposition took place in a coastal plain setting. however, during this period, fully marine conditions are only reported from the dutch part of the central graben (van adrichem boogaert & kouwe 1993). a coastline must therefore have been located in the southern part of the danish sector or in the german sector of the central graben. the regional transport direction in the danish central graben was probably towards the coastline in the south, parallel to the basin axis. within the middle jurassic succession, several levels of well-developed, erosionally-based channels, prograding deltas and splays, and lacustrine and floodplain mudstone units form correlatable units over a wide area, comprising the søgne basin and the tail end graben. this suggests a common, external control on base level, such as rift tectonics, eustatic sea-level change, climate change or a combination of these factors. in the søgne basin, the northern tail end graben and possibly in the salt dome province, the deposits of the aalen-1 – bat-1 sequences are cut by a significant erosion surface, the basal sequence boundary of cal-1 (fig. 7). the development of this surface was caused by a major fall in relative sea level. at the time of formation of this surface, the regional slope had changed from a southwards to a northwards dip (andsbjerg 2003, this volume). the erosional surface, which is commonly developed at the base of extensive fluvial or estuarine channel sandstones, is suggested to be the bounding 277 sequences lithostratigraphy chronostratigraphy lithostratigraphy chronostratigraphy sequencesgr sonic res kimm-2 kimm-1 baj-1 aalen-1 4874 m b. msl. lo la f or m at io n br yn e fo rm at io n lu lu fm m id dl e ju ra ss ic amalie-1 cal-1 u pp er ju ra ss ic up pe rm os t a al en ia n ba jo ci an ba th on ia n c al lo vi an u pp er k im m er id gi an o xf . – l. k im m . gr sonic res cal-1 bat-1 baj-1 3488 m b. msl. lo la fm lu lu fm m id dl e ju ra ss ic up pe rm os t a al en ia n ba jo ci an west lulu-1 aalen-1 u . ju r. t ri as si c u . k im m . ba th . c al lo vi an – l. k im m . br yn e fo rm at io n ba ct on g p bat-1 oxf-1 ox-1 – kimm-1 50 m maximum flooding surface flooding surface sequence boundary fig. 6. middle–late jurassic sequences in the amalie-1 and west lulu-1 wells illustrating, in particular, the characteristics of the middle jurassic (aalen-1 – cal-1) sequences. bath., bathonian; jur., jurassic; kimm., kimmeridgian; l., lower; oxf., oxfordian; u., upper. 278 k im m -1 o x2 o x1 c al -1 ba t1 ba j-1 a al en -1 c al -1 – k im m -1 ba j-1 a al en -1 20 k m a m al ie -1 w .l ul u3 lu lu -1 1 n or a1 m ar in e m ud st on e m ar in e si lts to ne m ar in e an d pa ra lic sa nd st on e fl uv ia l s an ds to ne fl oo dp la in m ud st on e an d si lts to ne m ax im um fl oo di ng s ur fa ce se qu en ce b ou nd ar y g r so ni c 38 31 m r es w . l ul u3 g r so ni c 38 41 m r es w . l ul u1 g r so ni c 36 99 m r es lu lu -1 g r so ni c 53 43 m r es a m al ie -1 g r so ni c 53 38 m r es n or a1 10 0 m fi g. 7 . se q u en ce s tr at ig ra p h ic c o rr el at io n o f th e m id d le an d l o w er m o st u p p er j u ra ss ic i n t h e sø gn e b as in a n d t h e n o rt h er n t ai l e n d g ra b en . a n u n co n fo rm ity a t th e b as e o f th e c al -1 s eq u en ce t ru n ca te s u n d er ly in g se q u en ce s. surface of an incised valley (andsbjerg 2003, this volume). the valley-fill is dated to the latest late bathonian – early callovian. in the søgne basin, the channel sands were initially deposited in straight or sinuous rivers, which show an increase in tidal influence with time. in the salt dome province, channel sandstones are either stacked or isolated and may show fining-upwards trends typical of sinuous channel fills or blocky gamma-ray motifs, indicating deposition in a straight or braided river (e.g. john flank-1, u-1; fig. 4; koch 1983). the incised valley-fill of the søgne basin is capped by a coal bed up to 3 m thick. later in the callovian, a low-energy coastal plain or delta plain characterised by small distributary channels, lagoons and coal swamps, was established in the southern part of the study area (koch 1983). contemporaneously, a barrier coast separated tidal lagoons at the margins of the søgne basin and the northern tail end graben from a wave-dominated marine bay in the central and eastern parts of these basins (fig. 9b). 279 ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■■ ■ ■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■■ ■ ■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■■300–400 m 200–300 m 100–200 m 50–100 m 0–50 m sequence absent fault well ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■■ ■ ■ ■■ ■ ■ ■■ ■■ ■■ ■■ a cal-1 ox-1 c ox-2 b 20 km 20 km 20 km fig. 8. isochore maps of the cal-1 (a), ox-1 (b) and ox-2 (c) sequences. the distribution of the cal-1 and ox-1 sequences illustrates a pronounced depocentre in the northern tail end graben, continuing south into the rosa basin. a second depocentre is present in the southern tail end graben – eastern salt dome province. the three maps chart a progressive transgression in the area of the ravn and elly wells (fig. 1), in the south-eastern part of the heno plateau. phase 4. rift pulse related transgression (early oxfordian – early kimmeridgian) sequence definitions ox-1 sequence this sequence is restricted to the tail end graben, the deeper parts of the søgne basin, the salt dome province and the flanks of the heno plateau (fig. 8b). the lower boundary of the sequence is located at the top of a well-developed coarsening-upwards interval (e.g. deep gorm-1; nora-1; fig. 10). at the western margin of the basin, a sandy lst up to 12 m thick may be present, erosionally overlying pre-jurassic deposits (e.g. elly-2; fig. 10). in most wells, the sequence consists of marine mudstones; only in elly-2 are sandstones a major component of the sequence. the thickness of the sequence attains a maximum of 189 m in nora-1. a significant depocentre occurs in the northern part of the tail end graben (fig. 8b). the lod of ctenidodinium continuum in the middle part of the sequence in u-1 indicates an early oxfordian age for at least the lower part of the sequence. combined with the lod of rigeaudella aemula close to the upper boundary in the falk-1 well, this indicates an early–middle oxfordian age for the sequence (fig. 3c). 280 56 ° 30 ’ 4° 00’ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■■ ■ ■ ■■ ■ ■ ■■ ■■ ■■ ■■ 56 ° 00 ’ 4° 00’ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■■ ■ ■ ■■ ■ ■ ■■ ■■ ■■ ■■ shelf shoreface/shore lagoon and barrier bayhead delta floodplain with river lake no deposition/no preservation fault well ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ a aalenian – bathonian/ earliest callovian early–middle callovian b c late callovian – earliest oxfordian 20 km 20 km 20 km fig. 9. palaeogeographic maps for the middle – earliest late jurassic. the maps illustrate the gradual transition from a terrestrially-dominated environment (a) through a tidallyinfluenced coastal plain (b), to a fully marine environment (c). deposits of the lowermost part of the middle jurassic succession may have been present in the western part of the danish central graben, and removed by erosion during hanging-wall uplift in association with early half-graben subsidence. due to lack of data, the distribution of middle jurassic deposits in the feda graben is uncertain. ox-2 sequence the distribution of this sequence corresponds to that of ox-1 (fig. 8c). the lower sequence boundary is normally located at the top of a coarsening-upwards interval (e.g. nora-1; fig. 10). in a few wells situated at the western margin of the basin, the lower sequence boundary may be located at the sharp base of an up to 10 m thick sandstone unit referred to the lst (e.g. elly-2). with the exception of these sandstones at the basin margin, the sequence consists of marine mudstones. the thickness of the sequence attains a maximum of 117 m in nora-1. a depocentre for this sequence extends from the northern tail end graben to the rosa basin (fig. 8c). the lod of rigeaudella aemula immediately below the lower sequence boundary in falk-1 combined with the lod of compositosphaeridium polonicum close to the upper boundary of this sequence in u-1 indicates an earliest late oxfordian age for this sequence (fig. 3c). kimm-1 sequence the sequence is missing from the north-western part of the study area (fig. 11a). the lower boundary of the sequence is normally located at the top of a coarsening-upwards interval (e.g. nora-1; fig. 10). in the heno plateau area, the mfs is normally situated immediately above a marked gamma/sonic spike (e.g. elly-2, ravn-1; figs 10, 12). 281 3373 m b. msl. cal-1 pliens-2 ox -1 ox-2 kimm-1 kimm-2 4226 m b. msl. kimm-2 ox-2 ox-1 kimm-1 cal-1 u pp er ju ra ss ic t ri as si c * * middle oxfordian bryne fm * * middle oxfordian deep gorm-1 elly-2 gr ressequences lithostratigraphy chronostratigraphy gr sonic ressequences lithostratigraphy chronostratigraphy gr sonic ressequences lithostratigraphy chronostratigraphy ox-1 3825 m b. msl. kimm-1 ox-2 kimm-2 lo la f or m at io n m . g ra b. fm fa rs un d fm fj er ri ts le v fm c al lo vi an l. o xf . u pp er o xf or di an u . o xf or d. – l. k im m . u pp er k im m . pl ie ns ba ch ia n u pp er ju ra ss ic m id dl e ju ra ss ic lo w er ju ra ss ic lo la f m h en o fm lo w er o xf . u . o xf .– l. k im . u pp er o xf . u pp er k im m . fa rs un d fm u pp er ju ra ss ic u pp er o xf or di an – lo w er k im m er id gi an lo w er o xf or di an lo la f or m at io n c al lo vi an m id dl e ju ra ss ic m . g ra be n fm u pp er o xf or di an u pp er k im m er id gi an m id dl e o xf or di an nora-1 sonic maximum flooding surface flooding surface sequence boundary 50 m fig. 10. sequence stratigraphic breakdown of the lower upper jurassic in the deep gorm-1, elly-1 and nora-1 wells illustrating the features of the sequences ox-1 – kimm-1. kim./kimm., kimmeridgian; l., lower; m. grab., middle graben; oxf./oxford., oxfordian; u., upper. the sequence consists of marine mudstone. the maximum measured thickness is 250 m (nora-1). the sequence has a well-defined depocentre in the central part of the tail end graben (fig. 11a). the lod of compositumsphaeridium polonicum immediately below the lower boundary of this sequence in u-1 and the lod of scriniodinium crystallinum and of nannoceratopsis pellucida close to the upper boundary in a number of wells (e.g. amalie-1, anne-3, emma-1, john flank-1, m-8, u-1, west lulu-2) indicate a late oxfordian – early kimmeridgian age for this sequence (fig. 3c). basin evolution in the tail end graben, subsidence along the coffee soil fault accelerated in the early oxfordian resulting in a highly asymmetric distribution of the oxfordian – lower kimmeridgian sediment package and in very large thicknesses of sediments in the central and eastern parts of the tail end graben (figs 8a, 13; fig. 13 faces page 294). this suggests that the tail end graben had entered the rift climax stage. during this phase, fault-controlled sub282 a kimm-1 kimm-2 c kimm-3 b ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■■ ■ ■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■■ ■ ■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■■300–400 m 200–300 m 100–200 m 50–100 m 0–50 m sequence absent fault well ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ 20 km 20 km 20 km fig. 11. isochore maps of the kimm-1 (a), kimm-2 (b) and kimm-3 (c) sequences. comparison of a and b illustrates the progressive transgression of the northern heno plateau, the gertrud plateau and the feda graben area from the late oxfordian – earliest late kimmeridgian. subsidence of the southern feda graben began during the latest early – earliest late kimmeridgian. note the relatively uniform distribution of sediments within the northern part of the tail end graben. during deposition of sequence kimm-3 (late kimmeridgian), a central depocentre developed throughout the length of the tail end graben and the salt dome province, striking roughly nw–se. sidence in the danish central graben mainly occurred along n–s-trending faults (møller & rasmussen 2003, this volume). the lower oxfordian – lower kimmeridgian succession is characterised by an overall transgressive development during which the depositional environment shifted from paralic to fully marine. the marked transgression reflects the interaction of the increased subsidence rate (see above) and a eustatic rise in sea level that has been documented from the north sea rift system and elsewhere (hallam 1978, 1988; haq et al. 1988; surlyk 1990). fully marine conditions were established in the søgne basin in latest callovian – earliest oxfordian times (cal-1, tst) with the introduction of the shelf mudstones of the lola formation. a coastal plain environment probably still dominated parts of the salt dome province to the south, indicating that the callovian–oxfordian transgression entered the danish central graben from the north (figs 9b, c). during the early oxfordian, the tail end graben, the south-eastern marginal parts of the heno plateau, 283 4029 m b. msl. fa rs un d fo rm at io n kimm-2 u pp er ju ra ss ic u pp er k im m er id gi an kimm-1 ox-2 u pp er o xf or di an – lo w er k im m er id gi an kimm-3 lo la f or m at io n u pp er o xf or di an 5389 m b. msl. u pp er ju ra ss ic fa rs un d fo rm at io n kimm-3 kimm-2 u pp er k im m er id gi an h en o fm pe rm ia n z ec hs te in g ro up h en o fm ox-2 4026 m b. msl. u pp er ju ra ss ic u pp er o xf . kimm-2 kimm-1 u . o xf . – l. k im m . u pp er k im m er id gi an h en o fm lo la fm fa rs un d fm kimm-3 kimm-4 lo la f m kimm-4 kimm-4 nora-1 gr sonic ressequences lithostratigraphy chronostratigraphy gert-4 gr sonic ressequences lithostratigraphy chronostratigraphy ravn-1 gr sonic ressequences lithostratigraphy chronostratigraphy 50 m maximum flooding surface flooding surface sequence boundary fig. 12. upper jurassic stratigraphy of the gert-4, nora-1 and ravn-1 wells illustrating, in particular, the characteristics of the kimm-2 and kimm-3 sequences. kimm., kimmeridgian; l., lower; oxf., oxfordian; u., upper. the salt dome province, and the rosa basin were also transgressed (ox-1, ox-2; figs 8b, c, 9c, 13). the remainder of the heno plateau and possibly areas further west were slowly submerged during the late oxfordian – early kimmeridgian (kimm-1; figs 11a, 14; fig. 14 follows page 294). marginal marine sandstones were deposited locally on the southern part of the heno plateau during this transgression (e.g. elly-2; fig. 14) but marine mudstones of the lola formation were deposited over most of the basin. a number of minor transgressive–regressive cycles can be discerned within the overall transgressive succession of marine mudstones (e.g. nora-1; fig. 13). the main basinal part of the study area was thus characterised by an offshore environment during this period. an inner shelf to shoreface environment developed, however, on the gertrud plateau and the northern part of the heno plateau at the end of this depositional phase (fig. 15b). 284 ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■■shelf shoreface barrier and strandplain lagoon no deposition/preservation fault well ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■■ ■ ■ ■■ ■ ■ ■■ ■■ ■■ ■■ a late oxfordian – early kimmeridgian latest early – earliest late kimmeridgian c late kimmeridgian b 20 km 20 km 20 km fig. 15. palaeogeographic maps for the late oxfordian – late kimmeridgian (late jurassic). during the late oxfordian – early kimmeridgian (a), transgression continued across the northern heno plateau, the gertrud plateau and the feda graben areas (see also fig. 9c). after a major regression (base kimm-2 sb), marginal and shallow marine conditions dominated the plateau areas (b). subsequently, renewed transgression resulted in a westwards shift of the coastline and extension to the west of paralic and marginal marine conditions (c). phase 5. shallow marine deposits and changing structural patterns (late kimmeridgian) sequence definitions kimm-2 sequence this sequence is found throughout the study area east of the mads and inge highs (fig. 11b). the lower boundary has been placed at a conspicuous surface that, even in the deeper parts of the basin, shows an abrupt shift from high to low gamma-ray values, probably indicating an erosion surface (figs 10, 12). in parts of the study area, such as the feda graben, the northern part of the gertrud plateau and the northern part of the heno plateau, this surface coincides with the unconformity that separates the prejurassic and the upper jurassic deposits (e.g. gert-4; fig. 12). on the heno plateau, the lst is thought to be represented by a unit of conglomerates and coarse-grained sandstones less than 5 m thick (e.g. ravn-1; fig. 12) although it should be noted that johannessen (2003, this volume) refers this interval to the lowermost tst. in more distal settings, the lst is represented by a 10–15 m thick sandstone or sandy siltstone unit (e.g. elly-2; fig. 10). in the proximal settings of the heno and gertrud plateaus and the southern feda graben, the tst consists of paralic and shallow marine sandstones and mudstones. in basinal settings (e.g. nora-1), the tst is represented by a well-defined fining-upwards interval (fig. 12). this interval consists of silty claystone, siltstone or silty sandstone, becoming coarser grained towards the heno plateau. in the central part of the heno plateau (e.g. ravn-1), most of the hst was removed by erosion during the subsequent lowstand. on the northern heno plateau, the gertrud plateau and in the southern feda graben, the hst consists of a section that coarsens upwards from sandy siltstones to gravelly sandstones of shoreface origin (johannessen et al. 1996; johannessen 2003, this volume). in basinal settings, the hst consists of a coarsening-upwards interval of marine mudstones, siltstones and sandstones. the thickness of the sequence attains a maximum of 319 m in cleo-1. major depocentres for this sequence are located in the southern tail end graben – eastern salt dome province, in the northern tail end graben – søgne basin and in the feda graben (fig. 11b). a latest early kimmeridgian to earliest late kimmeridgian age is indicated by the lod of endoscrinium galeritum in the lower part of the sequence in edna-1 and west lulu-1 and the fad of subtilisphaera? paeminosa and s.? inaffecta in the lower part of the sequence in gert-1 combined with the lod in the sequence above of stephanelytron scarburghense in the gert-2, jeppe-1 and lulu-1 wells (fig. 3c). kimm-3 sequence the distribution of this sequence corresponds to that of the kimm-2 sequence (fig. 11c). in the central part of the heno plateau (e.g. ravn-1; fig. 12), the lower sequence boundary is represented by an erosion surface overlain by shallow marine conglomerates or gravelly sandstones. further north, on the heno plateau and in the gertrud plateau – feda graben area, the sequence boundary is marked by a thin conglomerate in several wells (johannessen 2003, this volume). in more basinal areas, south and east of the heno plateau, the lower sequence boundary is located at the top of a well-defined coarsening-upwards interval (e.g. nora-1; fig. 12). in the heno plateau – northernmost salt dome province area (e.g. ravn-1), the tst is developed as a backstepping set of parasequences that consist of marine silty sandstones, siltstones and claystones. in the feda graben – gertrud plateau area, the tst consists of shallow marine sandstones and siltstones that grade upwards to marine mudstones (e.g. gert-4; fig. 12). in some wells (e.g. jeppe-1; fig. 14), an abrupt change from marine sandstones to mudstones, interpreted as a flooding surface, replaces part of the coarsening-upwards section. in the tail end graben (e.g. nora-1), the tst is represented by a thin fining-upwards interval of marine mudstones. the hst typically consists of marine mudstones. in the feda graben – gertrud plateau area and in the søgne basin, the hst terminates in a distinct coarsening-upwards interval, which includes siltand sandstones (e.g. gert-4; fig. 12). the thickness of the sequence attains a maximum of 387 m in cleo-1. major depocentres are located in the søgne basin, the feda graben and the southern tail end graben – salt dome province (fig. 11c). the lod of stephanelytron scarburghense in the lower to middle part of this sequence in gert-2 and jeppe-1 and the lod of endoscrinium luridum in the succeeding sequence (kimm-4) indicate a late kimmeridgian age for this sequence (fig. 3c). basin evolution the basin configuration changed significantly during the late early kimmeridgian – late kimmeridgian. fault activity and fault-related subsidence waned and at the end of the period a new pattern of dominantly nnw– sse-trending faults was established (møller & rasmussen 2003, this volume). this development marked a pause between two rift pulses. at this time, the tail end graben, including the søgne basin and most of the heno plateau, formed one major half-graben with an eastwards-dipping hanging-wall slope. accommodation space had been created in the feda graben prior to this phase. the feda graben was separated from the tail end graben by a transfer zone (fig. 11b). the lack of rift-related subsidence, probably combined with a regional fall in sea level in the beginning of this phase, caused a significant relative sea-level fall. this sea-level fall caused the development of a distinctive 285 erosional sequence boundary traceable over most of the danish central graben (figs 13,14); it can be seen on seismic sections as an onlap surface (erik s. rasmussen, personal communication 1997). the succeeding transgression, caused by eustatic or regional relative sea-level rise, gave rise to a gradual flooding of the heno plateau, the southern part of the feda graben and the gertrud plateau area. two higherorder sea-level cycles can be recognised during this overall transgression; the kimm-2 sequence records the first cycle, the kimm-3 sequence the second. while marginal areas were characterised by deposition of sand in shallow marine to paralic environments, low-energy marine conditions prevailed in the deeper parts of the basin during the kimm-2 cycle (fig. 15b). sand was deposited in a high-energy shoreface environment on the southern part of the heno plateau and in a back-barrier environment on the northern part of the heno plateau and in the area of the gert ridge (johannessen et al. 1996; johannessen 2003, this volume). the flooding of the heno plateau area forced the coastline back from its position near the eastern margin of the plateau at the beginning of the transgression to the western margin of the heno plateau or possibly further west at the time of maximum flooding. subsequently, the shallow areas in the western part of the basin were separated from the marine basin to the east by a zone of sand-dominated prograding shoreface. the coastline was re-established in a position at the eastern edge of the heno plateau at the end of this cycle (fig. 15c). in latest early kimmeridgian time (lower kimm-2 sequence), the relatively limited accommodation space in the central parts of the heno plateau was rapidly filled with deposits of the lst and tst, before maximum flooding was attained (e.g. ravn-1, falk-1; fig. 14). a relatively thick tst was deposited in the feda graben due to a rapidly subsiding basin floor. after the time of maximum flooding, most remaining available accommodation space occurred in the tail end graben and the deeper parts of the feda graben where a thick hst was deposited (e.g. g-1, gert-4; fig. 14). by the end of the first sealevel cycle (kimm-2), an area consisting of the southern feda graben and the heno and gertrud plateaus was exposed or dominated by shallow-water conditions. deep-water conditions only prevailed in the central parts of the tail end graben and the søgne basin. the kimm-3 sequence is characterised by a major depocentre in the feda graben where thick marine deposits are related to the early onset of the next phase of subsidence. a central depocentre developed throughout the length of the tail end graben and the salt dome province, striking roughly nw–se (fig. 11c). coarse-grained conglomeratic shoreface deposits were deposited immediately above the sequence boundary (the base kimm-3 sb) on the heno plateau (johannessen 2003, this volume). these deposits were overlain by backstepping parasequences of shallow marine sand and silt (e.g. falk-1, ravn-1; figs 12, 13). the beginning of the next phase of subsidence caused a gradual change from coarseto fine-grained shallow marine deposits in basinal settings of the feda graben. an abrupt shift from shallow marine sandstone to offshore mudstone can be seen in marginal areas of the feda graben and on the gertrud plateau (e.g. gert-2, gwen-2, jeppe-1; fig. 14). when fully marine conditions were established in the feda graben, a very high rate of subsidence in this area impeded the re-establishment of paralic conditions during the highstand part of the cycle. in the tail end graben and søgne basin, the kimm-3 sequence consists of a thin unit of marine mudstones. at the beginning of kimm-3 time, accommodation space on the heno and gertrud plateaus was limited by an initial sea-level fall and a low rate of subsidence. the limited accommodation space available was mainly filled by deposits of the tst (fig. 14). in the remainder of the study area, there was sufficient accommodation space for the development of the hst (e.g. cleo-1, g-1; figs 16, 17, following page 294). on the southern part of the gertrud plateau and adjoining parts of the heno plateau, the remaining accommodation space was filled with shallow marine hst deposits followed by exposure and bypass (e.g. gwen-2; fig. 14). phase 6. deep-water mudstones in a composite graben: rift pulses and passive subsidence (latest late kimmeridgian – middle middle volgian) sequence definitions kimm-4 sequence this sequence has the same areal distribution as the kimm-2 and kimm-3 sequences (fig. 18a). on the heno plateau, the lower sequence boundary is typically placed where the lithology changes abruptly from the sandstones and siltstones of the underlying sequences to fully marine mudstones (e.g. ravn-1; fig. 12). in basinal areas, the lower sequence boundary is located at the top of a coarseningupwards interval, which in the feda graben (e.g. gert-4) and in the søgne basin (e.g. lulu-1; fig. 17) may be distinct. in the feda graben 286 – gertrud plateau area, the tst is well-developed (e.g. gert-2, gwen-2; fig. 19), whereas over much of the heno plateau, the salt dome province and in the tail end graben, it often occurs in a condensed form (e.g. nora-1; fig. 12). in general, the hst is much thicker than the tst (e.g. gert-2, nora-1; figs 12, 19); both the tst and hst consist of marine mudstones. the thickness of the sequence attains a maximum of 318 m in gert-4. the sequence has an elongated depocentre in the eastern part of the tail end graben and a secondary depocentre in the feda graben (fig. 18a). a late kimmeridgian age is indicated for the sequence based on the lod of stephanelytron scarburghense in the sequence below combined with the lod of endoscrinium luridum close to the upper sequence boundary in amalie-1 and cleo-1 (fig. 3c). volg-1 sequence this sequence has the same areal distribution as the previous sequences (fig. 18b). in the feda graben – gertrud plateau area and on the heno plateau, the lower sequence boundary is located at the top of a thin but conspicuous coarsening-upwards interval (fig. 19). in the more basinal settings of the tail end graben, the søgne basin, and the salt dome province, the boundary is situated at the turnaround point between rather indistinct coarsening-upwards and fining287 ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ n■■500–600 m 400–500 m 300–400 m 200–300 m 100–200 m 50–100 m 0–50 m sequence absent fault well ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■■ ■■ ■ ■ ■■ ■■ ■ ■ ■■■ ■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ a kimm-4 volg-1b c volg-2 20 km 20 km 20 km ■■ ■■ fig. 18. isochore maps of the kimm-4 (a), volg-1 (b), and volg-2 (c) sequences. during the latest kimmeridgian and earliest volgian (a, b), an asymmetric basin geometry was reestablished with an elongate depocentre near the coffee soil fault; a secondary depocentre developed in the eastern søgne basin during the earliest early volgian. the outer rough and ål basins were transgressed in the early volgian (c). upwards units (e.g. g-1, nora-1; fig. 16). in several wells in the salt dome province, the tst is missing and the mfs amalgamates with the underlying sequence boundary (e.g. deep gorm-1; fig. 14). the mfs is marked by a conspicuous peak on the gamma-ray, sonic and resistivity logs. both the tst and hst consist of marine mudstones. the thickness of the sequence attains a maximum of 210 m in gert-4. the main depocentre of the sequence is in the eastern part of the tail end graben and in the eastern søgne basin. a secondary depocentre is present in the feda graben (fig. 18b). the sequence is referred to the earliest early volgian (fig. 3c) based on the lod of endoscrinium luridum in the sequence below combined with the lod of subtilisphaera? paeminosa in the upper part of the sequence in a number of wells (e.g. amalie-1, deep gorm-1, gert-2, -4, gwen-2, ravn-2, u-1). 288 volg-4 kimm-4 4048 m b. msl. c ro m er k no ll g p c ro m er k no ll g p r ya z. r ya z. u pp er k im m . fa rs un d fo rm at io n volg-2 volg-1 volg-3 lo w er v ol gi an m .– u . v ol g. h en o fm volg-4 4018 m b. msl. * upper volgian m id dl e v ol gi an fa rs un d fo rm at io n volg-3 volg-2 volg-1 kimm-4 lo w er v ol gi an u pp er k im m er id gi an * 3930 m b. msl. fa rs un d fo rm at io n volg-3 u pp er ju ra ss ic u pp er k im m . m id dl e v ol gi an volg-2 volg-1 lo w er v ol gi an elin-1 gr sonic ressequences lithostratigraphy chronostratigraphy gert-2 gr sonic ressequences lithostratigraphy chronostratigraphy gwen-2 gr sonic ressequences lithostratigraphy chronostratigraphy u pp er ju ra ss ic l. c re t. l. c re t. u pp er ju ra ss ic 50 m maximum flooding surface sequence boundary fig. 19. upper kimmeridgian – volgian sequence stratigraphy of the elin-1, gert-2 and gwen-2 wells, highlighting the features of the kimm-4 – volg-3 sequences. kimm., kimmeridgian; l. cret., lower cretaceous; m.–u. volg., middle–upper volgian; ryaz., ryazanian. volg-2 sequence although partly or completely eroded in most wells in the salt dome province, this sequence is recognised throughout the remainder of the danish central graben, including the outer rough and ål basins (fig. 18c). the lower sequence boundary of this sequence is rather indistinct in most wells, being located at the top of a weak coarsening-upwards trend within the hst of the volg-1 sequence below (fig. 19). in basinal settings, where the sequence is relatively thick, the tst is four to five times thicker than the hst; in settings where the sequence is relatively thin, the tst and hst are of comparable thickness. the sequence consists of marine mudstones. the maximum thickness of the sequence measured in wells is 344 m (e-1) although seismic data indicate thicknesses in excess of 400 m in the prominent tail end graben depocentre. a secondary depocentre is present in the feda graben (fig. 18c). the lod of subtilisphaera? paeminosa in the sequence below (volg-1) and of oligospaeridium patulum in the sequence above (volg-3) indicate an early volgian age for this sequence. this is further supported by the lod of cribroperidinium? longicorne in the lower part of the sequence in the eg-1 and emma-1 wells (fig. 3c). volg-3 sequence the sequence is present over much of the study area but is missing locally in the south-western part of the salt dome province and in the area around the mandal high, probably due to erosion (fig. 20a). the lower sequence boundary is positioned at the top of a coarsening-upwards interval (e.g. elin-1, gert-2; fig.19). in a number of wells where the section is relatively thin, the sequence boundary is placed at the base of a somewhat coarser bed (e.g. gwen-2; fig. 19). the volg-3 sequence consists of marine mudstones. the thickness of the sequence attains a maximum of 259 m in e-1. the primary depocentre in the tail end graben branches into the arne–elin graben. secondary depocentres are present in the gertrud and feda grabens and in the outer rough basin (fig. 20a). the lod of oligosphaeridium patulum in the lower part of this sequence in a number of wells (e.g. deep gorm-1, elly-2, falk-1, gert-2, gert-4, i-1, m-8, u-1, v-1) combined with the lod of occisucysta balia in the middle to upper part of the sequence in the bo-1, gert-2, gwen-2 and ravn-2 wells indicate a latest early volgian – middle middle volgian age for this sequence. this is supported by the lod of perisseiasphaeridium pannosum in the upper part of the sequence in gert-1 and gert-2 (fig. 3c). basin evolution in this depositional phase, the occurrence of rift pulses is reflected by syn-rift successions with fining-upwards log patterns separated by successions with overall coarsening-upwards patterns, as described from the mesozoic of east greenland by surlyk (1978). a rift pulse in late kimmeridgian time is indicated by the asymmetric sed289 ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■■ ■■ ■ ■ ■■ ■■ ■ ■ ■■■ ■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■■ ■■ ■ ■ ■■ ■■ ■ ■ ■■■ ■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■200–300 m 100–200 m 50–100 m 0–50 m sequence absent fault well a volg-3 volg-4b 20 km 20 km fig. 20. isochore maps of the volg-3 (a) and volg-4 (b) sequences. in the latest early – middle middle volgian (a), subordinate depocentres developed in the arne–elin graben and in the outer rough basin. the asymmetric half-graben geometry of the tail end graben became less pronounced during the middle–late volgian (b). iment distribution of the kimm-4 deposits in the tail end graben (fig. 18a). another pulse in the earliest early volgian is reflected by fining-upwards log patterns in the lower part of the volg-1 sequence (e.g. g-1, gert-4; fig. 14) and by seismic data from the tail end graben (møller & rasmussen 2003, this volume). a further rift pulse in the early volgian is interpreted from the abrupt change from low to high gamma-ray values immediately below the volg-2 mfs (e.g. gert-4, nora-1; figs 14, 16). the latter two rift pulses are recorded on seismic data from the tail end graben where they resulted in a rotational tilt package (møller & rasmussen 2003, this volume, fig. 7). the tail end graben and the heno plateau acted initially as one major asymmetric basin, with an elongate, nw–se-oriented depocentre located in the tail end graben. during the early volgian, this depocentre expanded into the eastern part of the søgne basin, where subsidence increased significantly (fig. 18a, b). the feda graben continued as an important depocentre. the gertrud graben and the outer rough and ål basins became actively subsiding elements at the time 290 ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■shoreface sands offshore/basinal muds basin floor sands submarine fan no deposition/no preservation fault well ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■■ ■■ ■ ■ ■■ ■■ ■ ■ ■■■ ■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ a late kimmeridgian earliest early – late volgian b c late volgian – late ryazanian 20 km 20 km 20 km fig. 21. palaeogeographic maps for the late kimmeridgian – late ryazanian. in the early volgian, the outer rough and ål basins were transgressed (compare a and b). fan and basin floor sandstones were deposited in deep parts of the basin during the late volgian – ryazanian (c). of deposition of the volg-2 sequence (figs 17, 18c). later during this phase, increased fault activity resulted in the generation of several minor sub-basins, for example the arne–elin graben which became a distinct depocentre in the early middle volgian during deposition of the volg-3 sequence (figs 18c, 20a). plateau areas draped by heno formation sandstones were drowned and a coarsening-upwards succession of marine siltstones and mudstones of the farsund formation was deposited above the kimm-4 mfs (e.g. ravn-1; fig. 14). thin stormor gravity-derived sand layers were occasionally deposited in the søgne basin (figs 17, 21a). during the early volgian, the outer rough and ål basins were also transgressed, and fully marine conditions were established (figs 17, 21b). shoreface sands were deposited at the western margin of the outer rough basin, as documented by data from the uk sector of the central graben (mackertich 1996; spathopoulos et al. 2000). phase 7. organic-rich mudstones and basin axis turbidites (late middle volgian – early ryazanian) sequence definitions volg-4 sequence this sequence has not been found in the søgne basin and mandal high area; it is also missing from parts of the outer rough basin and the salt dome province. the sequence is present in all other parts of the study area (fig. 20b). the lower boundary of the volg-4 sequence is positioned at the top of the well-defined coarseningupwards hst of the volg-3 sequence below (e.g. elin-1; fig. 22). in a number of wells, the sequence boundary is further marked by the abrupt base of a slightly more silty or sandy unit (e.g. jeppe-1). in many wells (e.g. iris-1, jeppe-1), the lower part of the tst typically shows a fining-upwards/coarsening-upwards pattern; in v-1, this unit is developed as a 30 m thick sandstone unit (fig. 23). in these wells, the remainder of the tst consists of a short, distinct fining-upwards mudstone section. in other wells, the tst is represented by a simple fining-upwards interval of marine mudstones (e.g. bo-1, elin-1; fig. 22). the upper boundary of this sequence is commonly an erosion surface and the upper part of the hst is thus missing in many wells (e.g. gwen-2, jeppe-1; figs 19, 22). where a major part of the hst is preserved, it typically shows consistently high gammaray values (e.g. bo-1, elin-1). due to the increasingly organic-rich nature of the sediments (see below), the gamma-ray log is not considered a fully reliable grain-size indicator. despite the high gammaray readings, descriptions of the cuttings indicate the presence of sandand siltstone in this interval. it is therefore assumed that the high gamma-ray readings are caused by the presence of organic-rich ‘hot shales’ and not necessarily by a lack of sandand silt-sized material. the sequence consists of marine, organic-rich mudstone with siltand sandstone interbeds. the thickness of the sequence attains a maximum of 189 m in elin-1. the sequence shows a more even distribution than the previous sequences, but depocentres are still recognisable in the tail end graben, the arne–elin graben, the feda graben and the outer rough basin (fig. 20b). the lod of scriniodinium inritibile in the lower part of the sequence in bo-1 and of senoniasphaera jurassica in the lower to middle part of the sequence in a number of wells (e.g. bo-1, deep gorm-1, elin-1, gwen-2, i-1, jeppe-1, ravn-2, w-1) indicate a middle middle volgian age for the lower part of the sequence. in combination with the lod of egmontodinium polyplacophorum in the middle to upper part of the sequence in the bo-1, deep adda-1, e-1, i-1, lone-1 and v-1 wells, this indicates a middle middle volgian to late volgian age for the sequence. this age assumption is supported by the lod of one or more of the dinocyst species dichadogonyaulax pannea, glossodinium dimorphum, muderongia simplex (the form previously referred as muderongia sp. a; poulsen & riding 2003, this volume) and gochteodinia mutabilis within the sequence in a number of wells (e.g. bo-1, deep gorm-1, e-1, elin-1, gwen-2, i-1, iris-1, v-1; fig. 3c). ryaz-1 sequence the distribution of the ryaz-1 sequence corresponds to the distribution of the volg-4 sequence (see above). the development of the ‘hot shales’ of the bo member within this sequence and the transition to more calcareous sediments at the top of the farsund formation makes sequence stratigraphic analysis problematic. neither a mfs nor an upper boundary of the sequence have been picked. the sequence is not necessarily limited to the marine mudstones of the farsund formation but may straddle the boundary to the overlying cromer knoll group. the lower sequence boundary is normally distinct and often erosional (e.g. bo-1, elin-1, jeppe-1; fig. 22). turbidite sandstones occur above the boundary in several wells (e.g. iris-1, jeppe-1; fig. 23). in certain wells that lack such turbidite sandstones (e.g. e-1, gert-2), the bioevents of the turbidite-bearing section are absent; this hiatus is interpreted to have resulted from erosion or sediment bypass, updip from areas in which turbidite sands accumulated (ineson et al. 2003, this volume). the turbidite interval may represent lowstand and possibly early transgressive deposits. in areas without active turbidite transport or deposition, this interval is characterised by a gradual upwards increase in gamma-ray values, probably representing the background sedimentation of the basin (e.g. bo-1). this increase in the strength of the gammaray signal culminates in the highest gamma-ray values in the entire jurassic studied section. this interval corresponds to the ‘hot unit’ of jensen et al. (1986), formalised as the bo member (farsund formation) by michelsen et al. (2003, this volume). the dominant lithology is 291 organic-rich marine mudstone; the mudstones are locally interbedded with turbidite sandstones (ineson et al. 2003, this volume).the thickness of this interval, from the basal sequence boundary to the top of the farsund formation, attains a maximum of 326 m in iris-1. the lod of egmontodinium expiratum in the lower part of this sequence in a number of wells (e.g. bo-1, edna-1, gert-2) and of rotosphaerotopsis thula immediately above the ‘hottest’ part of the bo member (e.g. in anne-3, bo-1, e-1, edna-1, gert-2, jeppe-1) indicate a late volgian – early ryazanian age for this part of the sequence. the uppermost part of the farsund formation is of late ryazanian age. this is indicated by the lod of dingodinium spinosum which coincides broadly with the upper boundary of the farsund formation in a number of wells (e.g. deep gorm-1, e-1, edna-1, jens-1, jeppe-1, kim-1, lone-1, v-1; ineson et al. 2003, this volume) and by the lod of dichadogonyaulax culmula at the same stratigraphic level in the e-1, edna-1 and gert-2 wells (fig. 3c). basin evolution the rate of subsidence decreased in the tail end graben and the salt dome province during this phase; in combination with decreased sediment supply, this resulted in stratigraphic condensation. according to the time-scale of gradstein et al. (1994), the succession deposited during phase 7 represents a period of c. 5.5 million years, 292 4392 m b. msl. jeppe-1 * volg-2 volg-3 volg-4 ryaz-1r ya za ni an volg-3 2577 m b. msl. u pp er ju ra ss ic bo-1 ryaz-1 volg-4 m . v ol gi an fa rs un d fo rm at io n * * upper volgian * upper volgian gr sonic res 3741 m b. msl. c ro m er k no ll g p c ro m er k no ll g p c ro m er k no ll g p sequences lithostratigraphy chronostratigraphy gr sonic ressequences lithostratigraphy chronostratigraphy gr sonic ressequences lithostratigraphy chronostratigraphy elin-1 u pp er ju ra ss ic u . v ol g. volg-4 volg-3 m id dl e v ol gi an ryaz-1 fa rs un d fo rm at io n r ya za ni an lo w er c re ta ce ou s lo w er c re ta ce ou s 50 mmaximum flooding surface sequence boundary r ya za ni an lo w er c re ta ce ou s u pp er ju ra ss ic m id dl e v ol gi an fa rs un d fo rm at io n fig. 22. uppermost jurassic to lowermost cretaceous stratigraphy of the bo-1, elin-1 and jeppe-1 wells illustrating the characteristic features of the volg-4 – ryaz-1 sequences. m., middle; u. volg., upper volgian. 293 r ya z1 v ol g4 v ol g3 v ol g3 v ol g4 r ya z1 20 k m je pp e1 m on a1 ir is -1 el in -1 v -1 d ee p a dd a1 e1 m ar in e m ud st on e m ar in e si lts to ne m ar in e an d pa ra lic sa nd st on e m ax im um fl oo di ng s ur fa ce se qu en ce b ou nd ar y g r so ni c r es je pp e1 g r so ni c r es m on a1 g r so ni c r es ir is -1 g r so ni c r es el in -1 g r so ni c r es e1 g r so ni c r es d ee p a dd a1 g r so ni c r es v1 10 0 m46 45 m 42 02 m 45 50 m 41 30 m 33 48 m 29 79 m 31 03 m fi g. 2 3. s eq u en ce s tr at ig ra p h ic c o rr el at io n o f th e u p p er m o st u p p er j u ra ss ic a n d l o w er m o st c re ta ce o u s in t h e t ai l e n d g ra b en an d t h e g er tr u d g ra b en . fa n a n d b as in f lo o r sa n d st o n es a re co m m o n i n t h e u p p er v o lg ia n – r ya za n ia n s u cc es si o n . approximately equal to the time represented by the 2–10 times thicker succession of phase 6. fault activity ceased along large segments of the main boundary fault (møller & rasmussen 2003, this volume), and this caused the geometry of the tail end graben to change from an asymmetric rift to a more symmetrical saucer-like basin with a relatively uniform sediment distribution (fig. 20b). these changes are the first indications in the tail end graben of the beginning of an early post-rift stage. minor depocentres continued to exist in the northern part of the tail end graben and in the arne–elin graben. subsidence continued in the feda graben and in the outer rough basin. the deposits of phase 7 consist of marine mudstones of the farsund formation. they generally show a gradual increase in gamma-ray and toc values culminating in the ‘hot shales’ of the bo member in the uppermost part (figs 17, 23; jensen et al. 1986; ineson et al. 2003, this volume). below the bo member, a well-developed sequence boundary separates the volg-4 and ryaz-1 sequences. this sequence boundary is strongly erosional in most wells on the northern part of the heno plateau and along the southern margin of the feda and gertrud grabens (e.g. jeppe-1; fig. 17). above the sequence boundary, sandstones are present in a number of wells along the basin axes of the tail end graben and the gertrud graben (figs 21c, 23). the sandstones were deposited from turbidity currents (iris-1) and debris flows (jeppe-1) and may have originated from a number of sources, such as the east north sea block of the ringkøbing–fyn high, the mandal high and the gert ridge (nielsen 1985; ineson et al. 2003, this volume). their occurrence may be related to changed drainage patterns in the hinterland after subsidence had ceased along major segments of the main boundary fault, and to erosion of inverted structures or crests of rotated fault blocks. relative sea-level changes on the basis of the sequence stratigraphic framework presented here (fig. 24), a relative sea-level curve can be constructed for the middle–late jurassic time interval and compared with eustatic curves proposed by haq et al. (1988) and hallam (1988) based on north sea and global data and with the relative sea-level curve of surlyk (1990) for east greenland (fig. 25). the lower jurassic has only been penetrated by wells in a small part of the study area, and no attempt has been made to construct a sea-level curve for this part of the succession. data are lacking from the toarcian and most of the aalenian due to the regional uppermost aalenian or lowermost bajocian unconformity. good biostratigraphic datings become available from the uppermost bathonian–callovian and upwards; the sea-level curve thus has its starting point at this level. wells from the basin centre and intermediate positions were preferred in selecting data for the construction of the curve. the relative sea-level changes have been interpreted mainly from lithological changes (e.g. changes in mud content in marine sediments) reflecting bathymetrically related changes in energy level. lithological changes were interpreted from well log patterns and cores. sedimentary facies, palynofacies and the extent of marine flooding surfaces have also been used. high uranium contents in some marine mudstones, related to organic matter content rather than clay proportion, can cause large deflections on the well logs, thus hampering lithological interpretation. for that reason, sea-level interpretations were not attempted in the ‘hot shales’ of the bo member. the coastal plain deposits of latest bathonian – earliest callovian age were transgressed during the callovian – earliest oxfordian in most of the danish central graben. a relative sea-level rise caused a rapid, step-wise transgression that expanded the basin laterally and drowned previous sediment sources. the sea-level rise and transgression continued throughout the oxfordian and culminated in the early kimmeridgian. this development, from the latest bathonian to the late early kimmeridgian, is almost identical to the sea-level curves of haq et al. (1988) and hallam (1978, 1981, 1988), based on north sea and global data, and the curve of surlyk (1990) for east greenland. the gradual collapse of the north sea dome (hallam & sellwood 1976; ziegler 1982; underhill & partington 1993) and the initiation of rifting in the danish central graben may have influenced transgressive trends in the region during this period. however, the resemblance between the relative sea-level curve of this study and the curves of haq et al. (1988), hallam (1988) and surlyk (1990) may indicate that a common causal factor of super-regional or global extent also influenced sea-level changes during this interval. the uppermost lower kimmeridgian – upper kimmeridgian of the danish central graben is characterised by a double lowstand peak. the other sea-level curves are close to their maximum at this level. the lowstand in the danish central graben probably evolved as a result of a sea-level fall during a pause in rift-induced subsidence (phase 5). 294 after the lowstand event, the late kimmeridgian – late middle volgian interval shows an overall rise in relative sea level, interrupted by a number of minor falls. this overall trend differs from the fall seen in most of the other sea-level curves, probably due to the continued high subsidence rate in the danish central graben, which neutralised the effect of the global sea-level fall indicated by the other curves. the minor relative sealevel falls in the interval may reflect pauses between periods of active subsidence. lithology prediction an important reason for undertaking a sequence stratigraphic study of a hydrocarbon-producing basin is to improve lithology and reservoir prediction. the predictive potential in syn-rift units is much less than in post-rift units. this is due to the problems for orderly sediment dispersal caused by the creation of tilted fault blocks and sub-basins, by the continuous presence of accommodation space on the lower hanging-wall slopes of tilted fault blocks, and by the local supply of sand 295 fluvial and estuarine sandstones estuarine/lagoonal sandstones, heteroliths, mudstones and coal beds shoreface sandstones/siltstones offshore mudstones; locally sandy or silty deep marine sandstones organic-rich offshore mudstones floodplain/lacustrine mudstones hiatus sequence boundary chronostratigraphy se qu en ce st ra tig ra ph y ringkøbing– fyn high nesw danish central graben oxfordian kimmeridgian volgian ryazanian callovian bathonian bajocian l l l l l l u u u u u u m m m m aalenian a al en -1 – b aj -1 pl ie ns -1 , pl ie ns -2 ba t1 c al -1 1 2 r ya z1 k im m o xvo lg 1 1 2 2 3 3 4 4 toarcian pliensbachian 80 km fjerritslev fm bryne fm lulu fm heno fm lola fm farsund fm bo mb fig. 24. generalised sequence stratigraphic diagram of the middle jurassic – lowermost cretaceous succession. relative duration of stages after gradstein et al. (1994). l, lower; m, middle; u, upper. from erosion of uplifted footwall shoulders. however, in the middle to upper jurassic syn-rift deposits of the danish central graben, the sandstone units seem to show a systematic distribution, which may be related to the combination of sea-level changes and periods of little tectonic subsidence. the bathonian–callovian sandstones, that are widely distributed in the søgne basin, the tail end graben and the salt dome province, resulted from a generous supply of sand from the active denudation of the north sea dome area and a relatively low and only slowly increasing accommodation space in the rift initiation stage. the best reservoir sandstones occur in the lowstand and transgressive systems tracts of the baj-1, bat-1 and cal-1 sequences (figs 4, 7), in the form of laterally extensive fluvial channel sandstones and tidal and shoreface sandstones in the uppermost part of the sandstone unit, deposited prior to the final transgression. the desirable combination of large fluvial and tidal channels and extensive wave-reworked sandstones occurs in the deeper parts of the søgne basin and the tail end graben. in these areas, subsidence was sufficiently fast to create accommodation space for successions of wave-influenced sediments up to 30 m thick during the final middle jurassic transgression (andsbjerg 2003, this volume). an overall transgressive development is interpreted for the late jurassic. the largest concentration of reservoir sandstones in this succession (the heno formation) was deposited during a pause in the otherwise rapid subsidence. the sand is concentrated on the heno plateau, the gertrud plateau and along the southern margin of the feda graben. the heno plateau constituted an upper hanging-wall slope of a major half-graben, the tail end graben, and the gertrud plateau probably 296 valanginian age sea-level curveslow high ryazanian volgian kimmeridgian oxfordian callovian bathonian bajocian aalenian toarcian pliensbachian sinemurian hettangian short term long term dating uncertain dating uncertain danish central graben (this study) wollaston forland embayment jameson land embayment hallam (1988) surlyk (1990) haq et al. (1988) fig. 25. comparison of the relative sealevel curve of this study with the sealevel curves presented by haq et al. (1988), hallam (1988) and surlyk (1990). figure modified after surlyk (1990), according to the gradstein et al. (1994) time-scale. comprised a transfer zone between the tail end and feda grabens (figs 11, 13, 14). possible source areas for the sandstones were to the north-east, where the ula formation and bryne formation sandstones were located, and to the west, on the mid north sea high (figs 15b, c). fluvial systems and/or marine current and wave activity may have facilitated sediment transport to the gertrud and heno plateaus. a similar scenario was envisaged by howell et al. (1996) for the southern uk central graben. the turbidite and debris flow sandstones of the upper volgian – ryazanian constitute potential reservoirs (fig. 23). most of the sands were deposited during depositional phase 7, when subsidence had decreased, and parts of the danish central graben had broken up into fault-bounded sub-basins. inactivity of some coffee soil fault segments (møller & rasmussen 2003, this volume) may have increased the drainage area in the marginal hinterland and opened new sediment transport routes, causing an increased sand supply. sediment dispersal in the receiving basin was still hampered by a topography dominated by tilted fault blocks and many small subbasins. the most extensive sandstones are present immediately above the ryaz-1 basal sequence boundary. sandstones have not been encountered in the few released wells drilled in the danish part of the outer rough and ål basins. however, the presence of volgian reservoir sandstones of a type similar to the heno formation is reported from wells nearby in the uk sector (mackertich 1996; spathopoulos et al. 2000). conclusions the depositional history of the danish central graben in the jurassic shows a relationship to contemporaneous rift stages and to relative and eustatic sea-level changes. the basin development in the jurassic is subdivided into seven depositional phases that reflect this interplay: phase 1 (hettangian–pliensbachian). during phase 1, pre-rift marine mudstones were deposited in a stable epicontinental sea, that stretched across a major part of the north sea region. the deposits were removed from large parts of the central north sea region by subsequent erosion (phase 2). phase 2 (latest pliensbachian – latest aalenian/earliest bajocian). this corresponds to the tectonic protorift stage, being characterised by domal uplift and regionally extensive erosion. phase 3 (latest aalenian/earliest bajocian – late callovian). the middle jurassic sandstones owe their presence to erosion of the north sea dome, and to slow generation of accommodation space during initial rifting (initiation of the tectonic rift stage). subsidence rates increased in the callovian and oxfordian, heralding the onset of the climax of the tectonic rift stage. this resulted in marine flooding and cessation of sand deposition. phase 4 (early oxfordian – early kimmeridgian). in the oxfordian – early kimmeridgian, the tail end graben entered the rift climax resulting in the development of a half-graben. marine mudstones were deposited during this phase. phase 5 (late kimmeridgian). extensive sand deposition took place in the early late kimmeridgian during a halt in subsidence between two rift pulses. sand was deposited over a large area with low accommodation space, which was created on the hangingwall plateau. exposed highs and land areas outside the graben supplied the sand. due to the decrease in subsidence in the tail end graben, input of finegrained sediment could catch up with accommodation space generation, probably resulting in a flat topography at the end of this phase. phase 6 (latest late kimmeridgian – middle middle volgian). during this phase, renewed rifting caused rapid subsidence in the study area. several rift pulses occurred during this interval. due to the high overall rate of subsidence, sand deposition associated with relative sea-level falls did not occur. the danish central graben began to break up into minor subbasins during this phase. phase 7 (late middle volgian – ryazanian). subsidence decreased over large areas of the basin, which had probably entered an early post-rift stage. major segments of the coffee soil fault became inactive, resulting in a shallower, more symmetrical basin. active subsidence continued in a number of subbasins. highly organic-rich mudstones are prominent in the deposits of this phase. the mudstones are interbedded with turbidite and debris flow sand deposits locally, in particular along the tail end graben – gertrud graben basin axis and along the eastern margin of the tail end graben. 297 comparison of the sea-level curve constructed for the danish central graben with published sea-level curves indicates that deposition in the periodically very rapidly subsiding rift basin was strongly influenced by tectonics. in the latest bathonian – late early kimmeridgian, the overall deepening trend resulted from both eustatic sea-level rise and rift-related subsidence. during the rest of the late jurassic, the deepening trend was primarily related to the high subsidence rate associated with rifting. most sediments were deposited between rift pulses in accomodation space that was generated during the rift pulses or by passive subsidence; only a relatively minor proportion of the sediments was deposited during the rift pulses. acknowledgements we had fruitful discussions with colleagues jon r. ineson, peter n. johannessen, jens j. møller and erik s. rasmussen. we gratefully acknowledge the comments of jon r. ineson and erik s. rasmussen on an early manuscript and those of the two referees, m.j. fisher and d.n. parkinson. the work formed part of a ph.d. study at copenhagen university for the first author, who extends his gratitude to supervisor finn surlyk for his thorough and critical comments. eva b. koppelhus and niels e. poulsen kindly supplied us with palynological datings. the work was supported by the danish energy agency (efp-92, 1313/920002), mærsk oil and gas a/s and norsk hydro udforskning a.s./amerada hess a/s. references andersen, c., olsen, j.c., michelsen, o. & nygaard, e. 1982: structural outline and development. in: michelsen, o. 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(eds): tectonic evolution of the north sea rifts, 1–36. oxford: clarendon press. manuscript received 30 may 1997; revision accepted 9 february 2000. ryaz-1 volg-4 volg-3 volg-2 volg-1 ryaz-1 volg-4 volg-3 volg-2 volg-1 kimm-4 kimm-2 kimm-3 kimm-1 ox-2 ox-1 cal-1 bat-1 baj-1 aalen-1 20 km elin-1 falk-1 ravn-1 ravn-2 nora-1baron-2 marine mudstone marine siltstone marine and paralic sandstone fluvial sandstone floodplain mudstone and siltstone maximum flooding surface sequence boundary gr sonic 4401 m res ravn-2 gr sonic 4307 m res ravn-1 gr sonic 4174 m res falk-1 gr sonic 5233 m res baron-2 gr sonic 4719 m res elin-1 gr sonic 5338 m res nora-1 100 m fig. 13. sequence stratigraphic correlation of the middle and upper jurassic on the heno plateau and in the tail end graben; the profile represents a dip section from the heno plateau to the tail end graben. ryaz-1 volg-4 volg-3 volg-2 volg-1 kimm-4 kimm-3 kimm-2 kimm-1 volg-2 volg-1 kimm-4 kimm-3 kimm-2 100 m 20 km gert-4 gert-2 jeppe-1 ravn-1 falk-1 g-1 deep gorm-1 elly-2 gwen-2 marine mudstone marine siltstone marine and paralic sandstone fluvial sandstone floodplain mudstone and siltstone maximum flooding surface flooding surface sequence boundary gr sonic 5352 m res gert-4 gr sonic 4876 m res gert-2 gr sonic 5041 m res jeppe-1 gr sonic 4355 m res gwen-2 gr sonic 4307 m res ravn-1 gr sonic 4174 m res falk-1 gr sonic 4080 m res elly-2 gr sonic 3935 m res deep gorm-1 gr sonic 3816 m res g-1 fig. 14. sequence stratigraphic correlation of the upper jurassic succession from the feda graben across the heno plateau to the southern tail end graben. onlap of the base jurassic unconformity can be seen on the southern part of the heno plateau (i.e. from deep gorm-1 to ravn-2). the base kimm-2 sequence boundary exhibits a distinct log break, interpreted as an erosional surface that cuts into underlying marine mudstones in the southern part of the study area. the upper part of the kimm-3 sequence and possibly the lowermost part of the kimm-4 sequence onlaps the kimm-3 flooding surface in the feda graben and on the southern heno plateau. cleo-1 amalie-1 nora-1 e-1 g-1 volg-1 kimm-4 kimm-3 kimm-2 cal-1 – kimm-1 ryaz-1 volg-4 volg-3 volg-2 volg-1 kimm-4 kimm-3 kimm-2 kimm-1 volg-2 gr sonic 4622 m res gr sonic 5122 m res gr sonic 4805 m res gr sonic 4719 m res gr sonic 4087 m res gr sonic 3816 m res elin-1 20 km amalie-1 cleo-1 elin-1 nora-1 g-1 e-1 marine mudstone marine siltstone marine and paralic sandstone fluvial sandstone floodplain mudstone and siltstone maximum flooding surface sequence boundary 100 m fig. 16. sequence stratigraphic correlation of the upper jurassic succession in the tail end graben; the profile is parallel to the basin axis. volg-2 volg-1 kimm-4 kimm-3 kimm-2 cal-1 – kimm-1 bat-1 baj-1 volg-2 volg-3 volg-4 ryaz-1 20 km cleo-1 lulu-1 w. lulu-2 karl-1 jeppe-1 lone-1 kim-1 tordenskjold-1 marine mudstone marine siltstone marine and paralic sandstone fluvial sandstone floodplain mudstone and siltstone maximum flooding surface flooding surface sequence boundary gr sonic 3532 m res tordenskjold-1 gr sonic 4602 m res kim-1 gr sonic 3797 m res lone-1 gr sonic 5041 m res jeppe-1 gr sonic 4819 m res karl-1 gr sonic 3944 m res w. lulu-2 gr sonic 3699 m res lulu-1 gr sonic 4709 m res cleo-1 100 m fig. 17. sequence stratigraphic correlation of the upper jurassic in the northern part of the danish central graben. the profile is perpendicular to the main structural elements from the outer rough basin in the west to the søgne basin in the east. the sub-basins were transgressed in a stepwise manner from east to west from the middle jurassic to the earliest early volgian. geological survey of denmark and greenland bulletin 28, 2013, 37-40 37 evaluation of total groundwater abstraction from public waterworks in denmark using principal component analysis brian lyngby sørensen and rasmus rønde møller in denmark water abstraction data have been collected since the late 1970s. initially the purpose was to monitor and assess the groundwater resources available for future local water abstraction. for this reason, abstraction data were collected not only from waterworks, but also from irrigation, industry etc. today water abstraction data are used for several purposes, for instance in water -balance calculations to estimate the available resource to wetlands, streams and lakes or to calculate the flow of chemical substances in the water environment. the role of climatic changes in the future hydrological cycle is subject to increasing attention. apart from a small reserve of surface water, all drinking water in denmark comes from groundwater. when precipitation changes in the future the amount of groundwater available for abstraction will also change. hence, for reasons of security of supply and environmental impact, it is important to know the amount and trend of abstraction each year. at national level, it is a statutory objective to abstract groundwater in a way that does not obstruct the general water-environmental objectives outlined in the european union’s water framework directive (the european parliament and the council of the european union 2000). the purpose of this paper is to present a method to evaluate the errors in the overall national groundwater abstraction dataset and describe how to correct erroneous data. for the sake of overview the national data are typically presented as an overall sum in million cubic metres per year (e.g. thorling et al. 2012). public groundwater abstraction in denmark drinking water in denmark comes from approximately 2500 waterworks, abstracting about 400 million m3 of groundwater per year. there is a pronounced decentralised water supply structure with many small waterworks spread across the country. approximately 72% of the waterworks each abstract less than 0.1 million m3 water per year, amounting to a total of 56.5 million m3 per year. at the other end of the fig. 1. an example of a time series for a specific municipality before (a) and after (b) correction of the abstraction data. data from 2011 are included in the graph for clarity. © 2013 geus. geological survey of denmark and greenland bulletin 28, 37–40. open access: www.geus.dk/publications/bull m ill lo n m 3 o f g ro un dw at er a bs tr ac te d m ill lo n m 3 o f g ro un dw at er a bs tr ac te d 0 5 10 15 20 25 1990 2000 2010 year corrected 0 5 10 15 20 25 1990 2000 2010 year a b uncorrected 3838 scale, 3% of the waterworks each abstract more than 1 million m3 per year, totalling 154 million m3 per year. according to danish legislation it is mandatory for waterworks and other users abstracting groundwater to report the amount abstracted once a year to the municipalities. the municipalities check for mistyped data and forward them to the national danish database on geology, groundwater and drinking water (the jupiter database at the geological survey of denmark and greenland). municipal reform in 2007, a major municipal reform took place in denmark. thirteen former counties (amter) were replaced by five socalled regions and most municipalities (kommuner) were merged into fewer and larger units, resulting in a drop from 271 to 98 municipalities. as part of this reform the new municipalities took over the responsibility to manage the water resources including abstraction licensing. this involved transferring employees from the former counties, new distribution of responsibilities and introduction of new computer systems and new procedures; all of which influenced the overall quality of the abstraction data. for instance, the new municipalities were responsible for submitting the 2006 water abstraction data to jupiter, although they were not operative before 1 january 2007. data preparation the water abstraction data used in this study were extracted from the jupiter database for the period 1989–2010. based on the extracted data, a date table was compiled with the sum of groundwater abstraction per year within each municipality. a time series for each municipality was plotted and visually inspected. at municipality level, small year-toyear changes and thus a smooth curve are expected, because fig. 2. total water abstraction in denmark for uncorrected (a) and corrected (b) data. the dashed lines show varexp – the correlation between the pca score of the first primary component (pc1) and the input data, expressed in million m3 per year. problem cause action no data were reported at all from the municipality an expected average was calculated based on data from 1–2 years before and after the year with missing data. evidently missing data no data from one or more waterworks. typing errors double registration from one or more waterworks. typing errors an expected average was calculated for the individual waterworks, or in case of typing errors a more probable value was estimated. evidently too high amount quoted evident double registrations were subtracted from the sum. in case of typing errors a more probable value was estimated. table 1. typical problems associated with registration of water abstraction data no data other apparent error unidentified no action taken. uncorrected corrected 300 350 400 450 500 550 600 1990 1995 20052000 2010 year a 300 350 400 450 500 550 600 1990 1995 20052000 2010 year b m ill lo n m 3 o f g ro un dw at er a bs tr ac te d m ill lo n m 3 o f g ro un dw at er a bs tr ac te d 39 on average the waterworks abstract almost the same amount each year. after initial inspection, 22 municipalities with unexpected data pattern were selected for detailed examination. four types of main problems were identified (table 1); the causes for three of the types could be identified and relevant action taken. correction of abstraction data for a single municipality an example of a time series for a selected municipality is shown in fig. 1a. the water abstraction from a specific waterworks was erroneously reported three times in the years 2006–2008 and twice in the years 2009–2010. thus, the water abstraction in the municipality was overestimated by 16.7 and 7 million m3, respectively, in the two periods. with the extra registrations removed, the time series shows a behaviour similar to what is expected (fig. 1b). a similar inspection was made of the time series from the 21 other municipalities. finally, a new data table was compiled by merging the corrected data with the data from the uncorrected time series from the remaining 76 municipalities. principal component analysis and pearson’s correlation coefficient principal component analysis is a mathematical procedure introduced by pearson (1901) and widely used to visualise multivariate data by dimension reduction (garcia & filzmoser 2011). according to garcia & filzmoser, the main problems of multivariate data can be avoided by using the principal component analysis to transform “. . . the original variables into a smaller set of latent variables which are uncorrelated”. each new variable (principal component or pc) can then be interpreted independently. there are several ways to perform principal component analysis, some of which are described in wikipedia (2013). the method used here is singular value decomposition (svd) using the ‘prcomp’ function of the base package of r (r core team 2012). the time series for the individual municipalities were used as objects (rows) and the years were used as variables (columns). for each year the pearson’s correlation coefficient ρ between the scores of the first principal component (pc1) and the corrected and uncorrected datasets d, was calculated and expressed in terms of million m3 (varexp) using the formula: where t is the total national abstraction. the correlation was done using the default settings of the ‘cor’ function of r (r core team 2012). the magnitude of ρ shows the strength of the linear dependence between the score of pc1 and d. status of water abstraction and comparison of uncorrected and corrected data figure 2 shows the total groundwater abstraction from public waterworks in million m3 per year from 1989 to 2010 with uncorrected and corrected data. both diagrams show the pearson’s correlation coefficient expressed in million m3 (varexp, dashed lines), according to the formula above. the variance explained ranges between 90 and 98% of the total yearly water abstraction. the remaining 2–10% can be perceived as ‘noise’ in the sense that this part of the variance is due to errors, short-term but large extra deliveries of water, abrupt changes in water needs, new or closed down waterworks etc. before the municipal reform (the period from 1989 to 2005) the unexplained variance on average corresponds to 16 million m3 for the uncorrected data and 12.7 million m3 for the corrected data. the improvement of the explained variance by correcting the data is thus 3.3 million m3. after the reform (2006–2010) the unexplained variance on average corresponds to 45.3 million m3 for the uncorrected data and 20.8 million m3 for the corrected data, leading fig. 3. locally weighted average (loess) of uncorrected and corrected groundwater abstraction data. uncorrected (data point / loess) corrected (data point / loess) 300 350 400 450 500 550 600 1990 1995 20052000 2010 year m ill lo n m 3 o f g ro un dw at er a bs tr ac te d 4040 to an average improvement of 24.5 million m3 by correcting the data. because of the errors mentioned above the amount of groundwater abstracted in denmark by the waterworks is only known with some uncertainty. in fig. 3 a locally weighted regression (loess) is calculated for corrected and uncorrected abstraction data in order to yield a ‘best guess’ of the total water abstraction. the curves show an overall trend with a large decline in the first half of the 1990s when abstraction decreased c. 20% from c. 550 million m3 in 1990 to c. 460 million m3 in 1996. later, the abstraction dropped to just over 400 million m3 in 2005. from fig. 3 it is clear that when corrected data are used, the abstraction flattens out at around 400 million m3 per year from 2005 onwards. if uncorrected data are used the abstraction level seems to decrease even further to below 400 million m3 per year over the same period. therefore the interpretation of trends depends to a large degree on whether the data are corrected or not. the main reasons for the large decline after 1989 are adoption of new legislation, increased water taxes and water saving campaigns (stockmarr & thomsen 2006). conclusions after the municipal reform in 2007 water abstraction data reported to the jupiter database show increased levels of errors due to changes in the way data are treated and reported. this means that national trends and levels are blurred which can lead to misinterpretations. by carefully examining data from the individual waterworks, it is often possible to determine the causes of errors and thereby correct them. the combined use of pca and pearson’s correlation coefficient is a useful way to provide an overall check on how well the data are corrected. this study shows that after the municipal reform the improvement is on average equivalent to 24.5 million m3 or c. 6%. on regional and local scales the impact of erroneous data can be severe. the example in fig. 1 shows that the abstraction can be overestimated by a factor 2.5 if no action is taken to investigate and correct erroneous data. it is crucial to correct and improve such data before they are used in waterbalance calculations, hydrological modelling, abstraction licensing and projections of water use in denmark. references garcia, h. & filzmoser, p. 2011: multivariate statistical analysis using the r package chemometrics, 71 pp. vienna: vienna university of technology, department of statistics and probability theory. thorling, l., hansen, b., langtofte, c., brüsch, w., møller, r.r. & mielby, s. 2012: grundvandsovervågning 2012 – grundvand. status og udvikling 1989–2011. teknisk rapport, http://www.geus.dk/publications/grundvandsovervaagning/1989_2011.htm pearson, k. 1901: on lines and planes of closest fit to systems of points in space. philosophical magazine, series 6, 2, 559–572. r core team 2012: r: a language and environment for statistical computing. vienna: r foundation for statistical computing, http://www.rproject.org/ stockmarr, j. & thomsen, r. 2006: water supply in denmark. the danish action plan for promotion of eco-efficient technologies – danish lessons, 18 pp. copenhagen: miljøstyrelsen. the european parliament and the council of the european union 2000: establishing a framework for community action in the field of water policy. http://eur-lex.europa.eu/lexuriserv/lexuriserv. do?uri=celex:32000l0060:en:html wikipedia, the free encyclopedia 2013: principal component analysis. accessed 7 february 2013. http://en.wikipedia.org/wiki/principal_component_analysis authors’ addresses b.l.s. & r.r.m.* geological survey of denmark and greenland, lyseng allé 1, dk-8270 højbjerg, denmark. e-mail: bls@geus.dk * present address: horsens kommune, rådhustorvet 4, 8700 horsens, denmark. http://www.geus.dk/publications/grundvandsovervaagning/1989_2011.htm http://www.geus.dk/publications/grundvandsovervaagning/1989_2011.htm http://www.r-project.org/ http://www.r-project.org/ http://www.ecoinnovation.dk/nr/rdonlyres/e4d4bd37-82e9-413d-87d8-d6aecd6b7e79/0/vandforsyning_artikel.pdf http://eur-lex.europa.eu/lexuriserv/lexuriserv.do?uri=celex:32000l0060:en:html http://eur-lex.europa.eu/lexuriserv/lexuriserv.do?uri=celex:32000l0060:en:html http://en.wikipedia.org/wiki/principal_component_analysis http://en.wikipedia.org/wiki/principal_component_analysis mailto:tvp@geus.dk geological survey of denmark and greenland bulletin 28, 2013, 9-12 9 drowning of the miocene billund delta, jylland: land–sea fluctuations during a global warming event erik skovbjerg rasmussen, torsten utescher and karen dybkjær lower miocene strata from boreholes and, in particular, at outcrops in the lillebælt and limfjorden areas of jylland provide a natural laboratory for studying the drowning of a major delta system during a period of global warming. detailed studies of sedimentary structures, fossil algae, spores and pollen give information about depositional environments, local temperatures and precipitation. by comparing with the global climatic record from the same period, a detailed reconstruction of the flooding of a low-relief delta system can be made, with emphasis on the global warming after the glacial event mi1a. the local temperature increase following the mi1a event is estimated to be c. 5°c. the billund delta during the early miocene, a delta system prograded into the eastern part of the north sea (fig. 1), with a sediment source in present-day norway and central sweden. for more than 100 million years, the eastern north sea was a relatively deep basin, but due to tectonism associated with the alpine orogeny and the opening of the north atlantic both the hinterland and the marginal areas of the basin were uplifted during the latest oligocene – early miocene (e.g. knox et al. 2010). inversion tectonism in the norwegian–danish basin resulted in shallower water depths of c. 100 m in the eastern north sea. contemporaneous uplift of the norwegian mountains provided a high sediment supply sourcing a major delta system prograding southwards into the north sea region. the delta front was shaped by wave action, similar to the present-day nile and danube deltas (bhattacharya & giosan 2003), with a size comparable to the latter. spit and barrier systems formed east of the main delta lobes (fig. 1) due to the prevailing westerly winds in the north sea area at that time. from the geological record, it is seen that the delta system prograded as far south as present-day southern jylland (fig. 1a). drowning of the land the depositional environments of the billund delta system are reconstructed from sedimentological studies (figs 1, 2). during the maximum extent of the delta, most of present-day jylfig. 1. palaeogeographic reconstruction of the early miocene in the eastern part of the north sea basin. a: billund formation during maximum regression. b: klintinghoved formation during initial transgression (kolding fjord member). c: klintinghoved formation during maximum transgression. modifided from rasmussen et al. 2010. © 2013 geus. geological survey of denmark and greenland bulletin 28, 9–12. open access: www.geus.dk/publications/bull 100 km 100 km vonsild rønshoved hagenør hostrup gyldendal vester thorsted b c 100 km a jyllandjylland north seanorth sea lillebæltlillebælt 1010 land was covered by braided fluvial channels and flood-plain environments (fig. 1a). in the late stage of delta progradation, the delta plain was irregular due to topographic elements formed associated with falling sea level, i.e. down-stepping delta platforms and incisions, and due to minor inversion of the basin (e.g. rasmussen et al. 2010). during later flooding, most of the deposits were redistributed by the action of waves. mud and fine-grained sand were winnowed away and transported into the deeper sea, whereas coarse-grained sand and gravel were reworked and redeposited locally. therefore, a widespread lag of gravel caps the shoreface sand of the underlying billund formation (fig. 3a, c). intense wave action on the main delta system shed sand towards the east, resulting in the formation of spit and barrier systems in eastern jylland (fig. 1b); lagoons and tidal flats developed north of these systems. both shoreface and lagoonal deposits accumulated in a depositional setting dominated by waves. the shoreface sand is dominated by hummocky and swaley, cross-stratified sand, and the lagoonal mud is commonly intercalated with washover fans (fig. 3b). from the sedimentary record, it is seen that the initial drowning of the billund delta system was characterised by flooding and re-establishment of land in two phases (fig. 2), which only influenced the distal part (tens of kilometres) of the delta plain. transgressive deposits up to 15 m thick are found at localities around lillebælt, rønshoved and hagenør (fig. 2). at hagenør, two stacked lagoonal mud units separated by shoreface sand show that relatively stable barrier systems were established twice during the transgression (figs 2, 3b). during the main flooding, the low-relief delta plain (similar present-day delta systems have gradients of c. 1/20 000) was flooded relatively quickly, and only up to c. 1 m of transgressive sediments were deposited, e.g. at hostrup (figs 2, 3c). at this locality, the relatively rapid changes in depositional environment during the main transgression are seen from the coexistence of the marine trace fossil ophiomorpha and rootlets (schaubcylendrichnus; fig. 3c). from studies of global sea-level changes in the early miocene, it appears that sea-level variations in the order of 25 m occurred during the mi1a event (miller et al. 2005). this change in sea level resulted in rapid progradation during the sea-level fall. the succeeding flooding can be followed for c. 75 km towards the north-east in the danish area (fig. 1c). maximum flooding has not been documented at any outcrop in denmark, but the sedimentary succession at hostrup indicates that the shoreline was located not far from this locality at any time because the marine mud of the klintinghoved formation is strongly influenced by terrestrial matter throughout the section. most of the 25 m of sea-level change must be explained by the 15 m transgressive deposits found at rønshoved and hagenør (compaction can be ignored in sorted shoreface sand) in the southern part of the delta system where some relief was created during the sea-level fall and lowstand. additional 5 m of relief of the main delta calculated from 75 km of flooding gives a total of 20 m of transgressive deposits. as the sea-level changes were in the order of 25 m (miller et al. 2005), the remaining 5 m can be ascribed to more or less stable conditions for a period (aggradation of the system) or waning of the inversion tectonism. in the upper part of the hostrup and gyldendal sections, the fig. 2. correlation panel of the lillebælt area and limfjorden. depositional environments mud sand gravel lithology in borehole datum top billund offshore sand shoreface sand barrier island sand marine mud klintinghoved fm kolding fjord mb billund fmbillund fm washover flat sand and mud lagoonal clay 0 gyldendal 5 m hostrup 0 5 m 0 5 m hagenør 0 5 10 m rønshoved (east) psa nd c l si vonsild 100 90 80 m b.s.l.m b.s.l. gamma ray log vester thorsted gamma ray log psa nd c l si psa nd c l si psa nd c l si 170 160 150 140 sse nnw 11 increased intercalation of hummocky, cross-stratified sand layers indicates resumed progradation of the shoreline. climate and vegetation change changes in continental climate and vegetation in the study area around the mi1a glacial event and the subsequent drowning of the billund delta system were reconstructed from pollen and spores (fig. 4). although terrestrial signals are commonly diluted in marine strata due to transport and reworking, our data show close correlation with the evolution recorded in the marine part. climate reconstruction using the ‘coexistence approach’ of mosbrugger & utescher (1997) shows distinct cooling during the formation of the billund delta and the sealevel lowstand that culminated at c. 164 m in the vester thorsted well (fig. 4). this is followed by warming that coincides with sea-level rise and drowning of the delta system. the cooling event is characterised by declining winter temperatures, with a fall of c. 5°c mean temperature in the coldest month. summer temperatures were stable to begin with, but later increased during progressive transgression (hagenør sample). the palynomorph record points to humid conditions with over 1000 mm annual rainfall over the time span analysed. the precipitation was not equally distributed throughout the year but showed distinct seasonality (cf. monthly precipitation records, fig. 4). the monthly means indicate that cooling was accompanied by declining precipitation during the wettest season and hence to a lesser degree of seasonal rainfall. fig. 3. photographs of lithologies of the outcrops of the billund and klintinghoved formations. a: shoreface deposits from rønshoved. note the gravel layer, the base of which forms the lower part of the klintinghoved formation. b: lagoonal deposits from hagenør. note that sand-rich washover fans dominate the upper part. c: transgressive sand and marine mud of the klintinghoved formation at hostrup. the lag of gravel at the base of the klintinghoved formation is seen in the lower part of the photograph. both rootlets and marine trace fossils are found in the sand-rich deposits. o: ophiomorpha, sch: schaubcylindrichnus (rootlets). fig. 4. climate records and synthesised pollen diagram for depth range 149–173 m of vester thorsted borehole. results obtained from lagoonal silts deposited in the trangressive phase of sequence c are shown on top. the lagoonal silts were sampled in the nearby hagenør outcrop. cmt: cold month mean, mat: mean annual temperature, wmt: warm month mean, mpdry, mpwarm, mpwet: mean precipitation of the driest, warmest and wettest month. a b ca o b c lag gravel rootletslagoon shoreface lagoon washover flat lag gravel billund fm billund fmbillund f rootletslagoon shoreface lagoon washover flat rootletslagoon shoreface lagoon washover flat o o sch sch 25 cm25 cm o o billund fm billund fm lag gravellag gravel 150 155 160 165 170 175 -5 0 5 10 15 20 25 30 temperature (°c) d ep th (m ) 1: pinus 2: far distance/alt. elements 3: lacustrine 4: swamp forest 5: coastal shrub 6: alluvial wetland vegetation 7: deciduous forest 8: evergreen forest/shrub 9: conifer forest 10: zonal herbs 11: ferns 0 25 50 75 100 palynomorph associations (%) 3000 50 100 150 200 250 precipitation (mm) vester thorsted hagenør cmt mat wmt mpdry mpwarm mpwet 1212 in the warmer climate reconstructed higher up in the transgressive systems track (sample hagenør), the region received more precipitation in the warm season. the palynomorph record also allows us to follow the evolution of continental ecosystems triggered by climate and sea-level change. the synthesised pollen diagram (fig. 4) shows frequency variations of groups reflecting regional to local biomes. pollen groups 1 and 2 comprise the pine family that includes many prolific pollen producers. association 1 mainly comprises pine, but in the hagenør sample small percentages of cedar and cathaya are also seen. the latter, monotypic genus occurs today as an endemic relict in central china. other members of pinaceae such as abies, picea and larix (group 2), frequently referred to vegetation of higher altitudes, are not important in the spectra (<3.5%). pine pollen can be dispersed over long distances and are relatively enriched in open marine palynomorph spectra (e.g. larsson et al. 2011). the overall declining trend of pine pollen recorded in the vester thorsted samples (from 45% at the base to 25% at 160 m) traces the falling sea-level and increased terrigenous input during the formation of the delta. the high pine pollen value for hagenør (>50%) coincides with the rising sea-level of the transgressive phase. lacustrine group 3, comprising lacustrine plankton, pteridophyte and angiosperm limnophytes, reaches its highest values (c. 5%) during the cooling event (at 164 m) and probably marks the most proximal conditions, together with fern group 11 that is frequent at the levels immediately above. these components originate from the local vegetation found on unstable, alluvial deposits on the delta plain. two groups of palynomorphs (4 + 5) that probably represent the vegetation of coastal swamps attain very high proportions during delta formation and show a distinct increase from c. 40 to over 55% which allows us to trace the ongoing regression. low values at hagenør (c. 10%) indicate that coastal swamps drowned as sea level rose. the arboreal vegetation on the coastal swamps comprises taxodioids (taxodium, sciadopitys and other cupressaceae), with ericaceae and myricaceae representing the shrub layer. pollen from alluvial wetland vegetation (6) play a minor role in the spectra and indicate a lack of widespread riverine forests, whereas palynomorphs usually assigned to zonal vegetation (groups 7–10) are present with 10 to 15% as permanent background. the zonal vegetation was a diverse, mixed conifer / broadleaved forest with evergreen and deciduous oak, magnolia, chestnut, sweetgum, members of juglandaceae, hemlock, coast redwood and palms. it is noteworthy that climate changes recorded in the section affected its generic composition. during the cool phase, between 156 and 164 m, the forest almost completely lost its broadleaved evergreen species, and thus had a more temperate aspect. concluding remarks during a global warming in the early miocene when the local temperature increased by 5°c, the billund delta system was flooded. the displacement of the shoreline was in the order of 75 km and affected the coastal vegetation. up to 15 m of alternating lagoonal and shoreface deposits were formed in the distal portion of the delta evidencing some stability in the early phase due to antecedent relief and high sediment supply. the main low-relief delta plain was, however, flooded quickly and only left c. 1 m of transgressive deposits. yearly precipitation rates were in the order of 1000 mm; the warming was accompanied by changes in seasonal patterns of rainfall so that during warmer periods there is a tendency to more summer rain. acknowledgements the nature agency centres in ribe, ringkøbing and aarhus are thanked for financial support. t.u. thanks the german science foundation (dfg) for finacial support. references bhattacharya, j.p. & giosan, l. 2003: wave-influenced deltas: geomorphological implications for facies reconstruction. sedimentology 50, 187–210. knox, r. et al. 2010: cenozoic. in: doornenbal, j.c. & stevenson, a.g. (eds): petroleum geological atlas of the southern permian basin area, 210–223. houten, the netherlands: european association of geoscientists & engineers (eage) publications. larsson, l.m., dybkjær, k., rasmussen, e.s., piasecki, s., utescher, t. & vajda, v. 2011: miocene climate evolution of northern europe: a palynological investigation from denmark. palaeogeography, palaeoclimatology, palaeoecology 309, 161–175. miller, kg., kominz, m.a., browing, j.v., wright, j.d., mountain, g.s., katz, m.e., sugarman, p.j., cramer, b.s., christie-blick, n. & pekar, s.f. 2005: the phanerozoic record of global sea-level changes. science 310, 1293–1298. mosbrugger, v. & utescher, t. 1997: the coexistence approach: a method for quantitative reconstructions of tertiary terrestrial palaeoclimate data using plant fossils. palaeogeography, palaeoclimatology, palaeoecology 134, 61–86. rasmussen, e.s., dybkjær, k. & piasecki, s. 2010: lithostratigraphy of the upper oligocene – miocene succession in denmark. geological survey of denmark and greenland bulletin 22, 93 pp. authors’ addresses e.s.r. & k.d., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: esr@geus.dk t.u., steinmann institute, university of bonn, nußallee 8, 53115 bonn; senckenberg research institute / bik f, senckenberganlage 25, 60325 frankfurt/ main, germany. mailto:hl@geus.dk geological survey of denmark and greenland bulletin 31, 2014, 67-70 67 combining exploration and multivariate techniques to detect the bjørnesund west gold occurrence, southern west greenland denis martin schlatter and bo møller stensgaard gold exploration in the bjørnesund region has been carried out since the early 1990s, and gold was found in the central part of the bjørnesund east area by nunaoil and the geological survey of denmark and greenland (geus). records of stream sediment samples with elevated gold concentrations up to several hundred parts per billion led to the recognition that amphibolites in the central part of the bjørnesund east could be a promising target and work in 1996 led to the discovery of hydrothermally altered amphibolites with up to several hundred ppb gold. however, exploration work was limited to grassroots prospecting and none of the targets were drilled. the aim of new field work was to target areas in the bjørnesund supracrustal belt which mainly consists of amphibolites but also comprises significant proportions of diorite, anorthosite, leucogabbro, granitoid rocks and ultramafic to mafic rocks that occur as relatively thin slivers in the amphibolite. we tested if the targeted areas were favourable for gold mineralisation and investigated the relationship between the mineral potential mapping and the actual geology. here we demonstrate that based on older data we located new gold mineralising systems in the western part of the bjørnesund supracrustal belt, identified platinum-enriched mafic to ultramafic rocks and located new occurrences of corundum at amphibolite-anorthosite contacts (schlatter © 2014 geus. geological survey of denmark and greenland bulletin 31, 67–70. open access: www.geus.dk/publications/bull fig. 1. geological map of the bjørnesund west and east areas showing the location of the newly discovered gold occurrence in bjørnesund west. prior to field work by geus in 2009, no gold occurrences were reported here. black outlines indicate the sampled areas of bjørnesund west, regarding sampling density see fig. 3. camp 1 bjørnesund west62°54´n bjørnesund east 62°56´n camp 2 1 km 50°03´w layers comprising garnet-rich amphibolite extensive surface staining new gold occurrence, bjørnesund west gold occurrence, bjørnesund east, 80 ppb fig. 3b fig. 3a bjørnesund dolerite granitic aplite/pegmatite trondhjemite (2.84 ga) trondhjemite with amphibolite inclusion ttg-gneiss (2.87–2.89 ga) migmatised ttg-gneiss ttg-gneiss with pl, bt and ep agmatitic ttg-gneiss ttg-gneiss with hbl and bt anorthosite ultramafic to mafic rock amphibolite amphibolite lens (hosted in gneiss) quartz-dioritic gneiss (2.919 ga) quaternary not digitised 50°18´w greenland 6868 & stensgaard 2012). we also show how lithogeochemical studies were useful to define the main rocks types, chemostratigraphic relations and hydrothermal alteration of the newly discovered gold mineralisation. based on our study, we encourage the use of artificial neural network analysis and data interpretation prior to field work in greenland in areas where only relatively little geological and mineral exploration work has been conducted and where the field season is relatively short. geology of the bjørnesund area the bjørnesund supracrustal belt (keulen et al. 2010) is of mesoarchaean age, c. 50 km long and a few hundred metres to 3 km wide (fig. 1). amphibolites dated to 2947 ± 47 ma are bounded towards the north and south by 2920–2810 ma tonalite-trondhjemite-granodiorite (ttg) gneisses (keulen et al. 2010; kolb et al. 2013) that are interpreted to have been intruded into the amphibolites. sheets of leucogabbro, gabbro and anorthosite are interpreted to have been intruded into the amphibolites at about 2950 ma and were in turn intruded by quartz dioritic protoliths at 2920 ma (keulen et al. in press). finally, late granites were intruded into the sequence of quartz-diorite amphibolite and anorthositegabbro between 2860 and 2830 ma (kolb et al. 2013). the rocks were affected by f1 folding into an isoclinal synform and f2 folding at 2850–2830 ma with an e–w-trending fold axis, which is the most dominant structural feature in the bjørnesund area (keulen et al. in press). the f2 folding is associated with thrusting that caused shearing with only minor displacement. finally, f3 nnw–sse-trending folds have bent the regional foliation slightly. methods a total of 116 rock samples (900 g) and 56 stream sediment samples (200 g) were collected during the field work and used for geochemical and petrographic investigations. rock samples were crushed and milled by actlabs laboratory in ontario, canada, and stream sediment samples were sieved at geus using a 0.18 mm sieve. the fine-grained fraction was sent to actlabs for analysis. gold was analysed by instrumental neutron activation whereas major and trace elements were analysed using actlabs package ‘4lithoresearch’. the u-pb ages of zircon grains from four plutonic rock samples were also determined by keulen et al. (in press). an artificial neural network is a mathematical and computational structure simulating the human neural network (the brain). information, in the form of input data, which are presented to the network causes the network to learn and recognise patterns in the data. for instance, when a network is shown multiple datasets for known gold occurrences, it is able to identify and memorise possible patterns in the datasets associated with the occurrences which can be regarded as training points for the network. afterwards, when the network is shown datasets from areas without any known gold occurrences, the network applies what it has learned and looks for patterns in the datasets that are similar to those recognised for the known gold occurrences. in that way, areas can be classified and mapped according to how similar their data patterns are to the patterns from the known gold occurrences. these areas can be regarded as potential to host gold occurrences. mineral potential mapping artificial neural network analysis (stensgaard 2013) was used for mineral potential mapping (fig. 2) together with ni/mg ratios from analysed stream sediment samples. this led to the identification of portions of the western and eastern parts of the bjørnesund supracrustal belt as the most favourable for gold occurrences. areas with anomalously high levels of fe3+ were identified using aster satellite images, and correspond to ultramafic dunitic and pyroxenitic rocks (schlatter & stensgaard 2012). the most favourable areas, as identified from artificial neural network analysis, coincide with areas of stream sediments with elevated ni/mg ratios. based on these detailed analyses, the bjørnesund west and east areas were selected for field work with the aim to characterise the geological environment and evaluate the gold potential. a new detailed and geo-referenced digital geological map was compiled after the new field work (fig. 1) and a new interesting gold occurrence was found in a hydrothermally altered shear zone in the amphibolites of the bjørnesund bjø rne sun d 5 km 50°10´w 62°55´n 2.1 3.2 3.9 4.9 6.2 7.5 fig. 2. neural network analysis for gold favourable areas in the bjørnesund west and east areas. top 8.5% most favourable areas for gold in seven coloured intervals according to the neural network analysis of as, cs, rb, sb and u stream sediment geochemistry and lineaments. a geological map is placed beneath the neural network analysis result, which is transparent. 69 west area. this several tens-of-metres wide ne–sw-trending shear zone (figs 1, 3a) dips 80° se. it can be followed over several hundred metres along strike and contains a 50 cm yellow-brownish, rusty-stained amphibolite, which hosts parallel quartz-carbonate-feldspar veinlets. chip samples collected over this altered unit yielded 569 ppb au (fig. 3a), and alteration related to this gold occurrence was identified to be of the garnet-biotite-iron oxide-hydroxide type (kolb et al. 2013). these new findings, together with elevated gold concentrations in stream sediment and rock samples along the same ne–sw-trending shear zone, indicate that the bjørnesund supracrustal belt may host undiscovered gold occurrences and confirm that stream sediment sampling is a powerful exploration tool (fig. 3a). furthermore, nickel and platinum-group elements could also constitute a target as pentlandite was identified in an ultramafic rock sample with elevated concentrations of ni, cr, co and pge from bjørnesund east (fig. 1; schlatter & stensgaard 2012). chemostratigrapy and hydrothermal alteration based on lithogeochemical immobile-element-ratio classification (barrett & maclean 1994) seven types of amphibolite and three different types of other mafic to ultramafic rocks (high mg-cr-ni-co, ni-rich and high-ti-zr basalt) were distinguished (figs 1, 3b). the gold horizon was located at the contact of basalt a and basalt e (fig. 3b). amphibolites with elevated gold concentrations (less than 100 ppb) were also located in bjørnesund east (fig. 1, c. 2 km north-east of camp 2) where the ore horizon is also located at the contact of basalt a and basalt e (schlatter & stensgaard 2012). it appears that this basalt a – basalt e contact represents a good geochemical marker horizon in the bjørnesund supracrustal belt. changes were calculated for 35 basalt samples from the bjørnesund area using the single precursor approach (maclean & barrett 1993). the results show that the richest gold-bearing basalt with 569 ppb au from bjørnesund west (fig. 3b, encircled in white) is characterised by strong additions of feo and silica (fig. 4a) and by gain of cao and loss of k2o (fig. 4b). in contrast, a basalt sample from bjørnesund east with 80 ppb au (fig. 1) shows only a small loss of silica, no change of iron, gain of k2o and loss of cao (fig. 4). we conclude that favourable alteration associated with gold mineralisation is characterised by silicification combined with addition of feo and cao. discussion and conclusions mineral potential mapping was successful because it indicated areas containing gold-mineralised rocks and pinpointed unusual mafic to ultramafic rocks where elevated concentrations of nickel and platinum-group metals were subsequently identified. extensive rust zones were identified from aster satellite data, some of which correspond to the mafic to ultramafic rock units with elevated ni, cr, co and pge contents; several of the ultramafic to mafic rocks fall into the komatiite field and pentlandite was identified by microprobe analysis in one sample (schlatter & stensgaard 2012). fig. 3. a: gold anomaly plot of sediment and rock samples from the bjørnesund west area. a several tens-of-metres wide shear zone (dashed line) with gold-mineralised portions was discovered at 62°54.4´n, 50°16.2´w and at 555 m a.s.l. legend below figure. b: chemostratigraphic relation seen from the bjørnesund west area. a rock unit which comprises mainly basalt a and a unit comprising mainly basalt c can be identified based on lithogeochemical immobile element techniques. the gold occurrence is found in rocks of basalt e type, and have basalt a and c in the structural footwall. for legend see fig. 1. quaternary granite amphibolite amphibolite of extrusive origin diorite and tonalite trondhjemite with amphibolite inclusion gneiss ultramafic rocks anorthosite b 31 122 16 16 21 569 9 50 11 71 6 11 134 6 au content in ppb rock sample stream/scree sediment sample 1 km 1 km gabbro basalt a field of basalt a field of basalt c ni-rich basalt basalt d basalt d basalt d basalt c basalt c high ti-zr basalt basalt c basalt x basalt a basalt f basalt f basalt b basalt e basalt e au mineralisation basalt a ttg gneiss amphibolite-gneiss anorthosite high mg-cr-ni-co basalt high mg-cr-ni-co basalt cam p 1 bjørnesund westa 97 7070 chemostratigraphic interpretations show that the au horizon is located at the contact between basalt a and basalt e and can be followed along the bjørnesund supracrustal belt for at least 10 km (figs 1, 3) so that this horizon represents an exploration target for gold and provides evidence of an e–w continuation. alteration related to au mineralisation is of garnet-biotite-iron oxide-hydroxide type and quartz-carbonate-feldspar veinlets occur in an amphibolite-hosted thrustshear zone between quartz-dioritic gneiss and gneiss. masschange calculations show that favourable ‘au-alteration’ is characterised by gains of feo, sio2 and cao and enrichment of as, sb and zn (schlatter & stensgaard 2012). the spatial association of gold occurrences and granite-trondhjemite rocks (fig. 1) possibly indicates that the granitoids played a role in the emplacement of the gold. the association of gold and nickel is intriguing (schlatter & steensgaard 2013) and could be related to deep structures which might have been activated during events similar to those described by fiorentini et al. (2012) from the highly nickel-enriched agnew-wiluna greenstone belt in western australia where felsic and komatiitic magmas are related to deep crustal conduits. with respect to the timing of events of the bjørnesund supracrustal belt, we suggest that ultramafic rocks, leucogabbro and anorthosite were intruded into amphibolites (fig. 1). this event may have been coeval with orthomagmatic nickel-platinum group-element mineralisation and isoclinal f1 folding. it is suggested that gold was then orogenically emplaced between 2860 and 2830 ma (kolb et al. 2013) at peak regional metamorphism. finally, late f3 deformation events created faulting in a staircase-like outcrop pattern at 2710–2700 ma (keulen et al. in press). this study represents a coupled effort of artificial neural network analysis and targeted field exploration and shows that such an approach can be efficient and successful in identifying new exploration targets of gold, nickel-platinum group elements and corundum. references barrett, t.j. & maclean, w.h. 1994: chemostratigraphy and hydrothermal alteration in exploration for vhms deposits in greenstones and younger volcanic rocks. in: lentz, d.r. (ed.): alteration and alteration processes associated with ore-forming systems. st. john’s: geological association of canada. short course notes 11, 433–467. fiorentini, m., beresford, s., barley, m., duuring, p., bekker, a., rosengren, n., cas, r. & hronsky j. 2012: district to camp controls on the genesis of komatiite-hosted nickel sulfide deposits, agnew-wiluna greenstone belt, western australia: insights from the multiple sulfur isotopes. economic geology 107, 781–796. keulen, n., kokfelt, t.f. & scherstén, a. 2010: notes on the common legend to the 1:100 000 digital geological map of southern west and south-west greenland, 61°30´–64°n. danmarks og grønlands geologiske undersøgelse rapport 2010/119, 41 pp. keulen et al., in press: mesoand neoarchaean geological history of the bjørnesund supracrustal belt, southern west greenland: settings for gold enrichment and corundum formation. precambrian research. kolb, j., dziggel, a. &. schlatter, d.m. 2013: gold occurrences of the archean north atlantic craton, southwestern greenland: a comprehensive genetic model. ore geology reviews 54, 29–58. maclean, w.h. & barrett t.j. 1993: lithogeochemical techniques using immobile elements. journal of geochemical exploration 48, 109–133. schlatter, d.m. & stensgaard, b.m. 2012: evaluation of the mineral potential in the bjørnesund greenstone belt combining mineral potential mapping, field work and lithogeochemistry. danmarks og grønlands geologiske undersøgelse rapport 2012/60, 60 pp. schlatter, d.m. & stensgaard, b.m. 2013: evaluation of the mineral potential in the bjørnesund greenstone belt, southern west greenland, combining multivariate studies, field work and geochemistry. 11th swiss geoscience meeting, lausanne, 15–16 november 2013. abstracts, 74–75. stensgaard, b.m. 2013: analysis of regional data sets: predictive gold potential using neural network analysis. danmarks og grønlands geologiske undersøgelse rapport 2013/15, 27–81. authors’ addresses d.m.s.* & b.m.s., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. * present address: helvetica exploration services gmbh, carl-spitteler-strasse 100, ch-8053 zürich, switzerland. e-mail: denis.schlatter@helvetica-exploration.ch a b ∆cao (wt%)∆sio2 (wt%) ∆f eo (w t% ) –10 0 10 20 30 40 50 60 –20 –10–40 0 20 0 1040 high ti-zr basalt (n=4) chemical rock types: basalt a (n=10) basalt c (n=2) basalt d (n=5) basalt e (n=10) basalt f (n=4) silicification addition of feo and silica removal of k2o and cao gain of cao and loss of k2o addition of k2o and loss of cao addition of feo and loss of silica silica removed au=569 ppb au=97 ppb au=97 ppb au =5 69 p pb au=80 ppb au=80 ppb au=134 ppb au=134 ppb ∆k 2o (w t% ) –2 –1 1 2 0 fig. 4. hydrothermal alteration based on mass-change calculations for 35 rocks from the bjørnesund west and east areas. a: δsio2 versus δ feo. b: δcao versus δk2o. mass changes were calculated using the method described by barrett & maclean (1994). mass changes are reported in wt% change (δ) relative to the precursor rock. geological survey of denmark and greenland bulletin 1, 301-347 301 sedimentology and sequence stratigraphy of the bryne and lulu formations, middle jurassic, northern danish central graben jan andsbjerg the middle jurassic bryne and lulu formations of the søgne basin (northern part of the danish central graben) consist of fluvially-dominated coastal plain deposits, overlain by interfingering shoreface and back-barrier deposits. laterally continuous, mainly fining-upwards fluvial channel sandstones that locally show evidence for tidal influence dominate the alluvial/coastal plain deposits of the lower bryne formation. the sandstones are separated by units of fine-grained floodplain sediments that show a fining-upwards – coarsening-upwards pattern and locally grade into lacustrine mudstones. a regional unconformity that separates the lower bryne formation from the mainly estuarine upper bryne formation is defined by the strongly erosional base of a succession of stacked channel sandstones, interpreted as the fill of a system of incised valleys. most of the stacked channel sandstones show abundant mud laminae and flasers, and rare herringbone structures, suggesting that they were deposited in a tidal environment, probably an estuary. several tens of metres of the lower bryne formation may have been removed by erosion at this unconformity. the estuarine channel sandstone succession is capped by coal beds that attain a thickness of several metres in the western part of the søgne basin, but are thin and poorly developed in the central part of the basin. above the coal beds, the lulu formation is dominated by various types of tidally influenced paralic deposits in the western part of the basin and by coarsening-upwards shoreface and beach deposits in central parts. westwards-thickening wedges of paralic deposits interfinger with eastwards-thickening wedges of shallow marine deposits. the middle jurassic succession is subdivided into nine sequences. in the lower bryne formation, sequence boundaries are situated at the base of laterally continuous fluvial channel sandstones whereas maximum flooding surfaces are placed in laterally extensive floodplain or lacustrine mudstones. the unconformity that separates the alluvial plain deposits of the lower bryne formation from the estuary deposits of the upper bryne formation is interpreted as a sequence boundary that bounds a system of incised valleys in the western and southern parts of the basin. sequence boundaries in the lulu formation are situated at the top of progradational shoreface units or at the base of estuarine channels. maximum flooding surfaces are located within marine or lagoonal mudstone units. marine highstand deposits are partitioned seawards, in the eastern part of the basin, whereas paralic transgressive deposits are partitioned landwards, in the west. this marked sediment partitioning in the uppermost part of the succession resulted from the alternation of episodes of fault-induced half-graben subsidence with periods of slow uniform subsidence. keywords: danish central graben, middle jurassic, bryne formation, lulu formation, sedimentology, sequence stratigraphy, alluvial/coastal plain – shallow marine, sediment partitioning geological survey of denmark and greenland, geocenter copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ja@geus.dk geological survey of denmark and greenland bulletin 1, 301–347 (2003) © geus, 2003 during the middle jurassic, the north sea area was dominated by extensive coastal plain, delta plain and shallow marine environments. the resultant deposits have been described from the viking graben (graue et al. 1987), from the moray firth and the yorkshire coast along the western margin of the north sea basin (hancock & fisher 1981; rawson & wright 1995; stephen & davies 1998), from the norwegian–danish basin along the eastern margin (nielsen 2003, this volume), and from the central graben in the central and southern north sea (gatliff et al. 1994; herngreen et al. 2003, this volume; fig. 1). in the past two decades, several minor gas, condensate and oil fields with middle jurassic reservoirs have been discovered in the søgne basin, a minor sub-basin straddling the danish–norwegian boundary line along the eastern main boundary fault of the central graben. production from these fields has started recently. the aims of this paper are threefold: (1) to provide a detailed environmental interpretation of characteristic sedimentary facies of the middle jurassic rocks and establish their palaeogeographic relationships; (2) to establish a high resolution sequence stratigraphic framework for the middle jurassic succession in the søgne basin of the danish central graben; and (3) to describe and interpret the important reservoir rocks in the upper part of the middle jurassic succession, their complex inter-relationships and their relationship to surrounding rocks, and the processes that caused such complexities. regional setting and structural development the danish central graben forms part of the central graben (fig. 1), a complex n–s-trending mesozoic intracratonic rift basin. subsidence of the danish central graben was initiated in the triassic but was most active during the middle and late jurassic (møller 1986). the central graben separates the mid north sea high to the 302 ■ ■ ■■ ■ ■ ■■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■ ■ ■■ ■■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ 55°n 4°e 100 km structural high outer moray firth v ik in g g ra be n central graben ringkøbing– fyn high n g nl uk dk mid north sea high normal fault national border well ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■■■■■ ■■ ■ ■ amalie-1 3/7-4 lulita-1 lulu-1 west lulu 1 3 2 4 ■■ c offee soil fault 20 km a b 57°30' 4° ringkøbing–fyn high east north sea block sørvestlandet high ■ ■ ■ ■ ■ ■ ■■ ■■ ■ ■ ■ ■ ■ ■ søgne basin tail end g raben ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ fig. 1. a: outline map of the jurassic north sea rift system showing the danish sector of the central graben (blue) and the position of the map in fig. 1b (red outline). dk, denmark; g, germany; n, norway; nl, the netherlands; uk, united kingdom. b: map of the northern danish central graben showing the structural outline of the søgne basin (grey), straddling the danish–norwegian boundary, and the location of wells used in this study. west from the east north sea block of the ringkøbing– fyn high to the east (japsen et al. 2003, this volume). the development of the danish central graben was determined by differential subsidence of grabens along n–sand nw–se-trending faults. the søgne basin and tail end graben began to subside as separate half-grabens during the middle jurassic (gowers & sæbøe 1985; møller 1986). initiation of rift-associated subsidence was probably related to domal uplift and subsequent dome collapse in the north sea area (whiteman et al. 1975; eynon 1981; ziegler 1982, 1990; underhill & partington 1993). rotation probably began in the søgne basin in connection with boundary fault activity during the middle jurassic (gowers & sæbøe 1985; møller 1986; cartwright 1991; michelsen et al. 1992; korstgaard et al. 1993) although it has been suggested that no syndepositional rotation took place in the søgne basin until volgian time (sundsbø & megson 1993). according to mogensen et al. (1992), salt structures were generated in the søgne basin in the triassic. middle jurassic subsidence and faulting initiated the development of boundary fault salt pillows and up-dip salt structures in the southern søgne basin. stratigraphic framework, concepts and methodology lithostratigraphy middle jurassic sandstones with interbedded mudstones and coals were encountered by the lulu-1 well, the first exploration well in the danish part of the søgne basin (fig. 1). similar deposits encountered in the norwegian part of the central graben were included in the bryne formation, a formation erected by vollset & doré (1984). jensen et al. (1986) extended the bryne formation to the middle jurassic deposits of the northern part of the danish central graben, although referring similar, coeval deposits of the southern part of the 303 fig. 2. jurassic lithostratigraphy of the danish central graben, from michelsen et al. (2003, this volume). system series stage ju ra ss ic t ri as si c åsgard formation leek member bo member heno fm lola formation bryne formation danish central graben ryazanian volgian kimmeridgian oxfordian callovian bathonian bajocian aalenian toarcian pliensbachian sinemurian hettangian rhaetian norian winterton formation c re ta ce ou s valanginian fjerritslev formation vyl fm poul fmfarsund formation lo w er u pp er m id dl e lo w er u pp er u l u l l u u m l l m u m l u l u u l l u m l u u l u m l paralic and non-marine sandstones, siltstones, mudstones and coals marine mudstones and siltstones unconformity offshore organic-rich marine shales submarine fan sandstones and siltstones shallow marine sandstones and siltstones hiatus middle graben formation lulu formation danish central graben to the central graben group of nam & rgd (1980). in a re-evaluation of the lithostratigraphy of the danish jurassic, the upper part of the middle jurassic succession in the northern part of the danish central graben is referred to a new formation, the lulu formation (fig. 2; michelsen et al. 2003, this volume). most of the middle jurassic succession in the southern part of the danish central graben previously referred to the central graben group is now included in the bryne formation, although the middle graben formation, defined from the dutch sector (see herngreen et al. 2003, this volume), is retained in this area (fig. 2; michelsen et al. 2003, this volume). the boundary between the bryne formation and the lulu formation is placed at the base of the first major coal or its correlative interval of thin coals and coaly mudstones in the upper part of the middle jurassic succession. in the søgne basin, the middle jurassic succession unconformably overlies triassic and permian deposits and is either succeeded conformably by marine mudstones of the upper jurassic lola formation (jensen et al. 1986) or is overlain unconformably by cretaceous deposits on structural highs. in the wells of the søgne basin, the thickness of the middle jurassic succession varies from 130 to 300 m; the succession may be absent from the top of structural highs, and it may attain a somewhat larger thickness in deeper parts of the basin. based on detailed sedimentological analysis of cores, the middle jurassic of the lulu-1 well was interpreted as deltaic interdistributary bay deposits overlain by coastal sediments (frandsen 1986). koch (1983) interpreted middle jurassic deposits further south in the danish central graben as alluvial plain and delta plain deposits. damtoft et al. (1992) suggested a fluvial channel and floodplain environment for the bryne formation, and johannesen & andsbjerg (1993) interpreted the middle jurassic succession in the søgne basin as an alluvial plain succession overlain by tidal and shallow marine deposits. biostratigraphy stratigraphically useful microfossils are rare in the studied succession. the sparse biostratigraphic information available for this study comes from unpublished reports from the geological survey of denmark and greenland, reports from service companies, and the results of new investigations prepared for the sequence stratigraphic study by andsbjerg & dybkjær (2003, this volume). only palynomorphs were used for dating in that study and the events are presented mainly as last occurrence datum (lod) of dinoflagellate cyst species. the events used and their relation to boreal standard zones are presented in andsbjerg & dybkjær (2003, this volume, fig. 3). age-specific microfossils have not been found in the lower part of the bryne formation in the study area. however, the occurrence of the dinoflagellate cyst scriniocassis sp. at 3730 m in west lulu-1 indicates an aalenian or earliest bajocian age. in the middle and upper parts of the bryne formation, the occurrence of the dinoflagellate cysts ctenidodinium combazii, impletosphaeridium varispinosum and the lod of the pollen quadraeculina anelliformis suggest a broad late bajocian to callovian age. more specifically, the occurrence of ctenidodinium combazii at 3742 m in the middle part of the bryne formation in west lulu-3 suggests an age not older than late bajocian for that interval. the occurrence of impletosphaeridium varispinosum in the incised valley deposits of the upper part of the bryne formation at 3602 m in west lulu-1 and at 3705 m in west lulu-3 indicates a latest bathonian to early callovian age. the lod of the pollen quadraeculina anelliformis either immediately beneath or just above the base of the incised valley deposits in several wells (e.g. 3717 m in west lulu-3, 4479 m in lulita-1) supports a latest bathonian age for valley incision and the initiation of valley infilling. the lulu formation, which constitutes the upper part of the middle jurassic succession, is poorly dated. however, the occurrence of durotrigia filapicata in the uppermost lulu formation at 4430 m in lulita-1 suggests an age not younger than the late callovian, and the lod of the dinoflagellate cyst liesbergia scarburghensis in the lower part of the lola formation in several wells (andsbjerg & dybkjær 2003, this volume), indicates a late callovian – mid-oxfordian age for the final transgression of the søgne basin. with ages spanning at least a period from the early bajocian to the latest callovian, the bryne and lulu formations represent about 18 ma of deposition, according to the time-scale of gradstein et al. (1994). sequence stratigraphic concepts and nomenclature sequence stratigraphic principles and nomenclature in this study follow posamentier et al. (1988, 1992), posamentier & vail (1988), van wagoner et al. (1988, 1990) and hunt & tucker (1992, 1995). andsbjerg & dybkjær (2003, this volume) present the subdivision and naming of sequences that can be 304 traced throughout the danish central graben. the present study attempts a more detailed sequence stratigraphic subdivision based on key-surfaces and units that are traceable across the søgne basin. andsbjerg & dybkjær (2003, this volume) subdivided the middle jurassic section into four sequences – the aalen-1 (aalenian), baj-1, bath-1, and cal-1 sequences. in this higher resolution local study, this nomenclature is retained but further subdivided (fig. 3). thus, cal-1 of andsbjerg & dybkjær (2003, this volume) is divided into cal-1a, 305 upper paralic wedge lithological units lithological units lithostratigraphy sequence stratigraphy upper paralic wedge upper marine wedge lola fm lulu fm lst/tst lst/tst hst/fsst lst/tst lst/tst lst/tst lst/tst hst hst hst hst sb mfs mfs mfs sb sb sb sb sb sb bryne fm lower marine wedge incised valley fill channel sand c channel sand b2 channel sand b1 channel sand a middle paralic wedge lower paralic wedge middle paralic wedge lower paralic wedge incised valley fill channel sand c channel sand b2 channel sand b1 channel sand a mfs mfs sb cal-1b cal-1c cal-1a bat-1b bat-1a baj-1b baj-1a aalen-1b aalen-1a sb fluvial channel sandstones floodplain mudstones estuary channel sandstones lagoonal/tidal flat mudstones/ heteroliths coal shoreface/mouth bar sandstones shelf mudstones fig. 3. architecture, lithostratigraphy and sequence stratigraphic interpretation of the middle jurassic in the northern part of the danish central graben. the middle jurassic sequences defined in this study are referred to according to their gross age, i.e. callovian sequences are termed cal-1a, cal-1b etc. system tracts: lst, lowstand systems tract; tst, transgressive systems tract; hst, highstand systems tract; fsst, falling stage systems tract. key surfaces: sb, sequence boundary; mfs, maximum flooding surface. cal-1b and cal-1c. the sequences and their most important key surfaces are shown in figures 3 and 4. the hierarchical nature of sequence stratigraphy, i.e. the potential subdivision of larger sequences into a number of smaller sequences, reflects the fact that sequences represent the varying time-spans over which different combinations of causal factors may operate. influenced by a variety of factors such as glacio-eustacy, tectono-eustacy, tectonics of various scales, and climate, sequences form over time scales ranging from tens of thousands of years to hundreds of millions of years (see discussion in vail et al. 1977, van wagoner et al. 1990, miall 1997). whereas the middle jurassic sequences outlined by andsbjerg & dybkjær (2003, this volume) represent time-spans of 5–10 ma. which is consistent with the influence of intraplate stress (cloetingh 1988; hallam 1988; miall 1997), the present study identifies sequences with durations in the range 1–5 ma., which may indicate a stronger influence of local tectonics. key surfaces and systems tracts a systems tract is defined as a linkage of contemporaneous depositional systems defined by stratal geometry at bounding surfaces, position within the sequence, and internal stacking patterns (posamentier et al. 1988). sequences are subdivided into the lowstand systems tract (lst), the transgressive systems tract (tst), the highstand systems tract (hst) and the falling stage systems tract (fsst; alternatively termed the forced regressive systems tract by hunt & tucker 1992, 1995). the lowstand systems tract (lst) consists of deposits formed at the lowest relative sea-level stand, bounded below by the mainly subaerial sequence-bounding unconformity (sb) and above by the first transgressive surface (ts). the tst consists of a succession of backstepping parasequences; individual parasequences may exhibit a progradational pattern. the lower boundary of the tst is the first ts and the upper boundary is the maximum flooding surface (mfs). the hst is characterised by a progradational stacking pattern, which may be interrupted by subordinate transgressive events. the systems tract is bounded at the base by the mfs and at the top by the sb or by a regressive surface of marine erosion (rsme) if it is overlain by a falling stage systems tract (fsst). the fsst consists of the sediments deposited during falling sea level and is bounded by the rsme at the base and by the sb at the top. there is a direct link between sequence development and relative sea-level change in the marine and marginal marine realm. in upland settings, sea-level changes do not influence sequential development of deposition significantly. a more pronounced influence may be present in non-marine deposits of lowland settings, although it may be subordinate to other factors. the sporadic occurrence of tidal indicators in the non-marine deposits of the bryne formation suggests that deposition took place on the lower part of a coastal plain where sea-level changes may have exerted a strong influence on sedimentation patterns and sequence development. data and methodology data from 9 released wells penetrating the bryne formation in the søgne basin were used in the present study (fig. 1). a total of 875 m of core has been examined and described. graphic core logs were matched to gamma-ray (gr) and sonic logs, supplemented by density, neutron and resistivity logs, in order to gain an improved interpretation of the cored successions. the observed core-to-log relationships have been used in the interpretation of well logs from uncored intervals by extrapolating sedimentological interpretations of cores to the uncored sections. the well logs and sedimentological core logs formed the basis for the construction of cross-sections. well-to-well correlations of key surfaces and characteristic units form a framework that guide correlations of other units and form the basis for the construction of palaeogeographic maps. sedimentary facies and depositional environments approximately 875 m of slabbed cores were available for the description of sedimentary facies. facies descriptions include the registration of lithology, grain size, primary sedimentary structures and deformation structures including degree and type of bioturbation. a total of 30 facies are recognised (table 1) and are grouped into nine facies associations (1–9), each of which represents a specific sedimentary environment. non-marine deposits sediments interpreted as mainly non-marine dominate the lower and middle part of the bryne formation. they are grouped into four facies associations repre306 c al -1 c c al -1 b c al -1 a lo la fm lu lu fm u pp er br yn e fm lo w er br yn e fm ba t1b ba t1a ba j-1 b ba j-1 a a al en -1 b a al en -1 a c al -1 c c al -1 b c al -1 a ba t1a ba j-1 b ba j-1 a a al en -1 b a al en -1 a g r g r g r g r g r g r g r d t d t d t d t d t d t d t w es t lu lu -4 w es t lu lu -2 w es t lu lu -3 w es t lu lu -1 3/ 74 a b1b2c lu lit a1 a m al ie -1 d ep os iti on al e nv ir on m en ts sh el f m ud st on es pa ra lic /s ho re lin e se di m en ts c oa l in ci se d va lle y fil l fl oo dp la in /la cu st ri ne s ed im en ts fl uv ia l c ha nn el s an ds to ne s pr eju ra ss ic s tr at a k ey s ur fa ce s se qu en ce b ou nd ar y m ax im um fl oo di ng s ur fa ce n or m al fa ul t c or ed s ec tio n 50 m ■ ■ ■ ■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■■ ■■■■ ■■ a m al ie -1 3/ 74 lu lit a1 lu lu -1 w es t lu lu 1 3 2 4 ■■ ■■ ■■ sø gn e ba si n 10 k m 307 fi g. 4 . w el l lo g p an el o f th e m id d le j u ra ss ic i n t h e sø gn e b as in . t h e d at u m l in e is t h e b as e o f th e lo w er m o st l at er al ly p er si st en t co al s ea m d ef in in g th e b as e o f th e lu lu f o rm at io n ( co al s ea m r 1 o f p et er se n & a n d sb je rg 1 99 6) . t h e fl u vi al c h an n el s an d st o n e u n its a , b 1 , b 2 an d c in t h e lo w er b ry n e fo rm at io n a re i n d ic at ed ( se e lu lit a1, a m al ie -1 w el ls ). n o te t h e p ro m in en t tr u n ca tio n o f m ar ke rs i n t h e lo w er b ry n e fo rm at io n b y th e m aj o r in ci se d v al le y d ef in ed b y th e c al -1 a se q u en ce b o u n d ar y. d t , so n ic v el o ci ty l o g; g r , ga m m ara y lo g. 308 facies structureless and laminated siltstone and claystone interbedded siltstone and sandstone bioturbated siltstone and sandstone hcs-dominated sandstone scs-dominated sandstone trough and planar crossbedded sandstone horizontally laminated and planar cross-bedded sandstone conglomerate and pebbly sandstone poorly sorted, bioturbated muddy sandstone and heterolith horizontally laminated and current rippled sandstone structureless rooted sandstone fining-upwards cross-bedded sandstone with mud drapes fining-upwards interbedded mudstone and sandstone coarsening-upwards sandstone with abundant mud laminae coarsening-upwards cross-bedded sandstone with mud laminae fining-upwards heterolithic sandstone and mudstone 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 description structureless or mmto cm-scale interlaminated siltstone and claystone cm-scale interbedded siltstone and sandstone. sharp-based, normal graded sand laminae show parallel lamination and wave ripples bioturbated siltstone, very fine-grained sandstone beds with sharp bases may show wave or combined flow ripples, parallel lamination and hcs very fineand fine-grained sandstone with siltstone laminae. sharp-based sandstone beds with hcs and subordinate wave ripple lamination fine-grained sandstone, scs, low-angle planar cross-bedding and scour structures fineto coarse-grained sandstone, occasionally pebbly. trough cross-bedding, planar cross-bedding and subordinate current and wave ripples fineto coarse-grained sandstone. parallel lamination with low-angle erosion surfaces and low-angle planar cross-bedding clast-supported pebble and granule conglomerate, less common matrix-supported conglomerate and pebbly sandstone. clast-supported conglomerate may show crossbedding; conglomerate veneers on erosion surfaces poorly sorted sandstone with subordinate siltstone. soft-sediment deformation structures and bioturbation dominate, some wave and current ripples may occur erosionally based fining-upwards units of very fineto medium-grained sandstone. parallel or gently inclined lamination, current ripples, local soft-sediment deformation structures or high-angle cross-bedding various sandstones and heteroliths fully or partly homogenised by roots fining-upwards units of fineto coarse-grained sandstone. planar and trough crossbedding with ripple cross-laminated flaser and wavy bedding in upper parts of units. abundant clay laminae; clay clasts and coal debris, interbedded sandstone and mudstone may occur fineand very fine-grained sandstone and heteroliths with mud clasts. ripple crosslamination, parallel lamination, flaser, lenticular and wavy bedding, cross-bedding very fineto fine-grained sandstone. bioturbated with mud laminae and flasers, ripple cross-lamination very fineto medium-grained sandstone and heteroliths. cross-bedding, crosslamination, flaser bedding, mud laminae thinly interbedded sandstone, mudstone and heteroliths. ripple cross-lamination, parallel lamination, flaser bedding table 1. facies classification of the bryne and lulu formations 309 thickness beds < 50 cm beds max. 10 cm silt beds less than 3 m, sand beds up to 10 cm, rarely to 50 cm 10–50 cm beds 20–50 cm beds 20–50 cm beds in units up to 3 m 5–15 cm beds in units up to 50 cm conglomerate beds max. 10 cm, pebbly sandstone beds up to 30 cm 5–10 cm beds in units up to 1 m 5–20 cm beds 50 cm – 3 m 4–10 m fining-upwards units fining-upwards units typically 50 cm – 2.5 m coarsening-upwards units up to 2 m coarsening-upwards units up to 4 m units less than 1 m biogenic structures weak to moderate bioturbation: anconichnus isp., planolites isp. and teichichnus isp. moderate to intense bioturbation: teichichnus isp., thalassinoides isp., skolithos isp., planolites isp. weak bioturbation trace fossils are rare bioturbation in the most fine-grained intervals: diplocraterion isp. and skolithos isp. trace fossils are rare: roots and (?)skolithos isp. often thoroughly bioturbated: teichichnus isp., diplocraterion isp. moderate bioturbation, roots may occur thoroughly homogenised by roots moderate, rarely intense bioturbation: teichichnus isp. moderate to intense bioturbation moderate to intense bioturbation: teichichnus isp. generally moderate bioturbation: teichichnus isp., diplocraterion isp. moderate to intense bioturbation: common diplocraterion isp., planolites isp. interpretation offshore, fair-weather deposition and suspension fall-out after storms offshore, near storm wave base offshore – offshore transition, storm activity alternating with long periods dominated by fair-weather conditions offshore transition, storm and waning storm deposition lower and middle shoreface, above fair-weather wave base upper shoreface. rip channels, nearshore bars and troughs foreshore beach and breaker zone deposits. may represent a transgressive lag transgressive marine sandstone deposited below fair-weather wave base during rising sea level washover sediments beach ridge plain major tidal channel or active tidal inlet tidal creek or inactive major tidal channel tidal sand bar/flat proximal flood tidal delta or estuary sand bar tidal flat and distal flood tidal delta 310 facies coarsening-upwards/fining-upwards sandstone and heterolith structureless or laminated mudstone and bioturbated sandstone organic-rich rooted mudstone coal fining-upwards interbedded sandstone and mudstone sandstone, fining-upwards or no grain-size trend intraformational conglomerate fining-upwards thin-bedded or cross-bedded sandstone chaotically bedded sandstone sideritic siltstone and mudstone coarsening-upwards units of deformed siltstone and sandstone coarsening-upwards units of sharp-based sandstone and siltstone disturbed silty mudstone organic-rich laminated mudstone 17 18 19 20 21 22 23 24 25 26 27 28 29 30 description coarsening-upwards very fineto medium-grained sandstone units. planar crossbedding, parallel lamination, current and wave ripple cross-lamination and small-scale hcs/scs. commonly associated with fining-upwards channel units parallel-laminated or structureless mudstone with interbeds and laminae of sandstone. parallel lamination, wave ripples, flaser and lenticular bedding organic-rich mudstones with plant fragments, thin coals and with abundant rootlets sharp-based units, often fining-upwards, of cross-bedded and cross-laminated sandstone with abundant heterolithic beds and mud laminae. abundant coal or mud clasts locally sharp-based fining-upwards cross-bedded and cross-laminated sandstone. heterolithic sandstones may dominate upper part of units; some beds may have abundant coal and mud clasts. thick amalgamated units may show no overall grain-size trend matrixor clast-supported, pebble–cobble conglomerate, with sand matrix. angular mudor siltstone clasts. conglomerate beds at base of fining-upwards sandstone units are parallel-stratified or cross-bedded sharp-based fining-upwards sandstones. thin-bedded with current ripple crosslamination, parallel lamination or cross-bedding. intraformational clasts and coal fragments, soft-sediment deformation poorly sorted sandstone with deformed and overturned mud laminae. coal and mud clasts scattered throughout siltstone and mudstone with siderite bands and nodules, abundant plant remains and roots. indistinct patches of sandstone may occur stacked coarsening-upwards units of siltstone and sandstone. dominated by softsediment deformation structures with current, wave, and climbing ripple lamination in sandstone units, parallel and climbing ripple lamination and wavy and lenticular bedding in siltstone units. mudstone clasts and coal fragments locally abundant. thinner, sharp based fining-upwards sandstones with deformed cross-bedding may occur at top of coarsening-upwards intervals coarsening-upwards units of very fineto medium-grained sandstone with siltand mudstone. common parallel lamination, current ripple cross-lamination, root traces, soft-sediment deformation. base gradational to floodplain mudstones mudand siltstone, subordinate sandstone, coal debris. parallel lamination, sediments disturbed by roots, soft-sediment deformation and pedogenesis organic-rich mudstones with sand and silt laminae table 1 (continued). facies classification of the bryne and lulu formations 311 thickness units up to 5 m units less than 2 m less than 50 cm max. 5 m units max. 12 m units max. 8 m beds up to 75 cm units < 2 m beds typically 20–50 cm typically 0.5–2 m 2–5 m units. may be stacked in 10 m coarsening-upwards successions beds 10–50 cm, units up to 2 m typically < 1 m max. 8 m biogenic structures moderately bioturbated: teichichnus isp., diplocraterion isp. moderate to intense bioturbation by roots moderate to intense bioturbation by roots upper part of channel units may be bioturbated: diplocraterion isp., teichichnus isp. thoroughly bioturbated, mainly by roots interpretation bay-head delta/bay shoreface low energy outer estuary, estuary central basin or lagoon marsh or vegetated coastal swamp mire tidally influenced fluvial channel major fluvial channel channel lag deposits crevasse channel or minor fluvial channel channel margin deposits of fluvial channels abandoned channel fill lacustrine delta. stacked minor coarseningupwards units capped by channel sandstones may represent delta lobes of a larger lacustrine delta levee and crevasse spray floodplain fines lake and pond 312 cm 0 10 20 30 40 50 60 70 80 90 a west lulu-3 3749.4 m b west lulu-3 3751 m c west lulu-3 3751.8 m d west lulu-3 3753.3 m e west lulu-3 3754.1 m f west lulu-3 3756.4 m sb this page and facing page: fig. 5. core photographs of fluvial channel and floodplain facies (facies associations 1–4) and facies successions of the lower bryne formation (baj-1b, bat-1a, bat-1b sequences). consecutive core sections in this and subsequent core photographs are bracketed. a–f: selected intervals of the fluvial channel unit b2 (sequence baj-1b, lst/tst) in west lulu-3 (base lower right, top upper left; for location of core sections, see fig. 13). the basal sequence boundary (sb) overlying sequence baj-1a lies immediately beneath the lowermost core section (f). sections b–f illustrate the active channel fill (facies association 1), showing trough cross-bedded, ripple cross-laminated and structureless sandstones, with abundant coal and mudstone clasts. these are succeeded (a) by passive channel fill or floodplain deposits (facies associations 2, 4) comprising mudstones, interbedded mudstones and sandstones and thin coals associated with rootlets 313 (arrowed) and palaeosol mottling. g–k: selected intervals from the fluvial channel sandstone unit c (sequence bat-1a) in west lulu-2 (base lower right, top upper left; for location of core sections, see fig. 14). the channel base defining the sequence boundary (sb, see core section k) is overlain by the active channel fill (facies association 1) comprising structureless and cross-bedded sandstones with abundant mudstone clasts and coal fragments (i–k, lower point bar) succeeded by sandstone and mudstone heteroliths, disturbed in places by bioturbation (rootlets arrowed) and soil-forming processes (g, h; upper point bar). the channel fill is capped by coal (g; facies association 4). l: sandstone and heterolithic sandstones showing climbing ripple cross-lamination, representing a lacustrine or crevasse delta (facies association 2). west lulu-1, sequence bat-1b; for location of core section, see fig. 15. cm 0 10 20 30 40 50 60 70 80 90 g west lulu-2 3851.1 m l west lulu-1 3626 m h west lulu-2 3859.3 m i west lulu-2 3860 m j west lulu-2 3861.6 m k west lulu-2 3862.2 m sb senting fluvial channel fill, proximal floodplain, lake and distal floodplain, and vegetated floodplain. facies association 1: fluvial channel fill (facies 21–26) description. the fluvial channel fill association comprises erosionally based, up to 8 m thick, fining-upwards, channel units. the channel units are dominated by sandstone in the lower part and become heterolithic in the upper part (fig. 5b–f). a conglomerate of mudstone clasts may occur immediately above the erosional base. the most common facies of the channel fill association are trough and planar cross-bedded sandstone and ripple crosslaminated sandstone (fig. 5b, d). the common chaotically bedded sandstone facies is characterised by contorted bedding, soft sediment deformation and a chaotic texture with abundant plant debris and intraformational mudstone clasts in places (fig. 5e). in the upper part of the channel units, sandstone beds are interbedded with 10–30 cm thick heterolithic beds that may represent inclined heterolithic strata (thomas et al. 1987), a variant of epsilon cross-stratification characteristic of tidally influenced fluvial channels (smith 1987). mudstone laminae, double mud drapes and flaser bedding, abundant in the sandstone facies of some units, particularly in the upper part of the bryne formation, suggest occasional tidal influence in the river system. interpretation. fining-upwards channel units that can be correlated between most wells in the study area (figs 3, 4), represent laterally extensive channel sandstones deposited by laterally migrating, sinuous rivers. chaotic bedding may be the result of bank collapse and/or post-depositional collapse of stems and other plant material deposited behind obstacles in the channel. similar deposits have been described by alexander & gawthorpe (1993; their facies s4) and by guion et al. (1995) as part of their minor channel facies. the evidence of occasional tidal influence suggests deposition in a coastal plain environment. facies association 2: proximal floodplain (facies 27–29) description. the proximal floodplain association consists of interbedded sandstone, siltstone and mudstone (fig. 5l). the sandstones are generally less than 2 m thick, but may be amalgamated into units 4–5 m thick. the sandstone units may fine upwards, coarsen upwards or show no overall grain-size trends. primary structures include cross-bedding, current ripple lamination, climbing ripple lamination, wave ripple cross-lamination, parallel lamination and chaotic bedding with abundant mudstone and coal clasts. soft sediment deformation structures are common. siltstones and mudstones of this association are commonly structureless but may show deformation structures, parallel lamination and lenticular bedding. interpretation. the sandstones were deposited in small channels, as crevasse splays, on levees and as small lacustrine deltas. sandstone units that show bi-directional current ripples, mud flasers and abundant mud laminae were probably influenced by tidal processes during deposition in fluvial channels or distributaries. siltstones and mudstones are interpreted as waning flow deposits on levees and in small fluvial and crevasse channels or as the passive infill of abandoned channels. facies association 3: lake and distal floodplain (facies 28–30) description. mudstones and siltstones dominate the lake and distal floodplain association (fig. 5k). interbedded sandstones are not thicker than a few decimetres. mudstones and siltstones form units up to 5 m thick; these are most commonly structureless or show parallel lamination. the parallel lamination is faint and may appear irregular and slightly deformed. heterolithic units may show lenticular and wavy bedding and current and wave ripples in thin sand beds. interpretation. the sediments are interpreted as having been deposited in ponds, shallow lakes and on the distal levee, or represent the passive infill of abandoned channels. facies association 4: vegetated floodplain (facies 20, 29, 30) description. sediments with abundant root traces, mottled siltstones and mudstones and coal beds are combined in this facies association. mottled siltstones and mudstones frequently have a light-coloured ‘leached’ appearance (fig. 5g). interpretation. the depositional environment was a floodplain where primary deposits were modified by 314 vegetation and soil-forming processes. most soils formed under reducing conditions. marginal marine deposits back-barrier and estuarine deposits dominate the upper part of the bryne formation and, in the western part of the basin, the lulu formation. the marginal marine deposits are separated into five facies associations (5–9), representing estuary channels and bars, flood tidal deltas and washover fans, bay-head deltas and bay-fill, lowenergy estuary and lagoon, and marsh and swamp. facies association 5: estuary channel and bar (facies 12–15) description. the estuary channel and bar association is represented by 4–10 m thick sandstone-dominated units that may fine upwards, coarsen upwards or show no clear grain-size trend, and as 0.5–4 m thick, finingupwards, fineto very fine-grained sandstones and heteroliths. the sandstones show planar and trough cross-bedding with common mudstone laminae (fig. 6d), and ripple cross-lamination with abundant mudstone flasers (figs 6c, e, 7). coal fragments and mudstone clasts occur in some beds, most commonly above erosional surfaces (fig. 6k, l). heterolithic strata, 5–20 cm thick, occur interbedded with the sandstones and may represent beds of inclined heterolithic stratification (thomas et al. 1987). interbedded sandstones, mudstones and heteroliths in the fine-grained units show flaser, wavy and lenticular bedding and parallel lamination (figs 6a, 8b, e). up to 5 m thick coarseningupwards units are formed by progressively thicker and coarser grained sandstone beds separated by thin mudstone and siltstone beds (fig. 9). these sandstone beds may show cross-bedding, ripple cross-lamination, and mudstone flasers and laminae, but may also be structureless with the exception of a few inclined mudstone laminae and mudstone flasers. intense bioturbation with abundant teichichnus isp. is common (fig. 9a, b). interpretation. the fining-upwards sandstone units, dominated by cross-bedding and ripple cross-lamination with abundant mudstone laminae, double mud drapes and flaser bedding, are interpreted as estuary point bar deposits (reineck & wunderlich 1968; visser 1980). sandstone units that show similar sedimentary structures but lack overall grain-size trends are interpreted as amalgamated tidal channel sandstones. the finer-grained fining-upwards units represent the passive infill of major channels or the active fill of minor channels. the coarsening-upwards sandstone units represent estuarine channel bars (fenies & tastet 1998) or mouth bar deposits of bay-head deltas. facies association 6: flood tidal delta and washover fan (facies 14–16) description. the flood tidal delta and washover fan association consists of up to 3 m thick, generally coarsening-upwards units of sandstones and heteroliths, that may be overlain by fining-upwards units of well-sorted sandstone (fig. 10f–h). in the coarsening-upwards units, fine-grained heterolithic beds show lenticular and wavy bedding. the most fine-grained sandstones are commonly strongly bioturbated, but some may show flaser bedding and parallel lamination. coarser grained sandstone facies include trough cross-bedded, parallellaminated and ripple cross-laminated sandstone. the coarser grained sandstones may have erosional surfaces overlain by thin conglomerates. the fining-upwards units are dominated by well-sorted, fineor very finegrained sandstone showing parallel lamination, lowangle planar cross-bedding, ripple cross-lamination and soft-sediment deformation structures. interpretation. the fine-grained heterolithic beds characterised by wavy and lenticular bedding and the coarsening-upwards sandstones with abundant mudstone laminae and flaser bedding, frequently interbedded with lagoonal mudstones, are interpreted as the deposits of flood tidal deltas and tidal sand flats. the coarsening-upwards trend and the association of physical structures correspond well with descriptions of recent flood tidal deltas (e.g. nichol & boyd 1993). the interbedding with lagoonal sediments further supports this interpretation. the well-sorted, erosionally based, finingupwards units that locally overlie flood tidal delta and tidal flat deposits show a close likeness to washover deposits described by schwartz (1982). facies association 7: bay-head delta and bay-fill (facies 4–7, 17, 18, 21) description. this association is typified by sandstones and heteroliths arranged in overall coarsening-upwards successions up to 8 m thick (fig. 9d–g). the sediments 315 316 cm 0 10 20 30 40 50 60 70 80 90 a west lulu-3 3667.3 m b west lulu-3 3669.9 m c west lulu-3 3670.8 m d west lulu-3 3671.7 m e west lulu-3 3674.2 m f west lulu-3 3680.2 m g west lulu-3 3684.6 m h west lulu-3 3685.4 m this page and facing page: fig. 6. core photographs of the incised valley-fill (facies associations 5, 8) of the upper bryne formation (cal-1a sequence) in west lulu-3 (base of succession lower right, top upper left; for location of core sections, see fig. 16). the erosional base of the incised valley, defining the cal-1a sequence boundary (sb), is observed in the lowermost core section (l) succeeded by the lower unit of active fluvial or estuary channel fills (j–l; facies association 5); this unit is dominated by well-sorted sandstone showing faint cross-bedding or chaotic bedding with abundant coal and mudstone clasts and a basal mudstone clast conglomerate immediately overlying the sequence boundary (l). lagoonal deposits (i; facies association 8) cap the lower channel unit, represented by burrowed mudstones showing signs of soil-forming processes, and are succeeded by inferred bay-head delta deposits (h). the upper unit of active (b–g) and passive (a) estuary channel fills (facies association 5) is characterised by sandstones with abundant double mud drapes (examples arrowed), flaser lamination (c) and cross-bedding (f, h). are dominated by current-generated structures, but wave-generated structures also occur. sandstones with current-generated structures may occur as channel deposits in the upper part of coarsening-upwards successions. minor units of well-sorted sandstone and heterolith may show low-angle planar cross-bedding, swaley and hummocky cross-stratification, and wave ripple cross-lamination. levels showing moderate bioturbation with teichichnus isp. are evident in places. deposits of this association frequently overlie fine-grained lagoonal deposits. interpretation. this association is interpreted to record the progradation of bay-head deltas into estuaries, lagoons, or bays. depending on the amount of wave influence, the deposits were either slightly modified by small-scale wave activity, or reworked thoroughly by storm wave activity. deposits may be difficult to distinguish from coarsening-upwards estuary bar deposits of association 5. facies association 8: low-energy estuary and lagoon (facies 14, 16, 18, 19) description. this facies association is represented by organic-rich mudstones, siltstones and heteroliths, showing parallel lamination, wavy and lenticular bedding, and ripple cross-lamination with mud-flasers, partly obliterated by biogenic activity (figs 8d, e, 9i). sedimentary units of this association vary in thickness from a few decimetres to several metres. interpretation. the dominance of finer grain sizes suggests deposition in a low-energy environment. the assemblage of sedimentary structures is typical of a tidally influenced environment such as an estuary central basin or a lagoon with extensive tidal flats. facies association 9: marsh and swamp (facies 19, 20) description. coals, mudstones and associated rooted heteroliths are grouped in the marsh and swamp association. mudstones and rooted heteroliths have a dark grey to black appearance, reflecting the high organic content (fig. 9c, d). both vitrinite-rich and inertiniterich coals are present. interpretation. the vitrinite-rich coals represent deposition in a waterlogged, anoxic mire environment. the 317 cm 0 10 20 30 40 50 60 70 80 90 i west lulu-3 3689.5 m j west lulu-3 3699.6 m k west lulu-3 3702.2 m l west lulu-3 3710.5 m sb inertinite-rich coals represent a somewhat drier environment, with periodically oxic conditions in a swamp or raised bog. pyrite in some coal beds suggests the occasional influx of marine water. the evidence of marine influxes and the association of the coals and rooted sediments with lagoonal deposits suggest that deposition took place in back-barrier swamps and marshes (petersen & andsbjerg 1996). marine deposits marine deposits dominate the lulu formation in the central parts of the søgne basin, but thin units can be traced into the mainly paralic deposits in the western part of the basin. the marine deposits are separated into three facies associations: offshore, prograding shoreface and beach, and transgressive shelf and shoreface. facies association 10: offshore (facies 1, 2, 9) description. the offshore association consists of up to 50 cm thick units of structureless and laminated mudstone, and cm-scale interbedded, heterolithic mudstone and sandstone. the association frequently forms coarsening-upwards units with structureless mudstone in the basal part overlain by heterolithic mudstone with siltstone and sandstone laminae and beds that show an upwards increase in thickness, grading into the more sandy deposits of the shoreface association (figs 11a, f, 12i). laminae may be normally graded, and show parallel lamination and wave and combined flow ripple lamination. the sandstone beds are commonly sharp-based. the sandstone-dominated upper part of coarseningupwards units may grade into hummocky cross-stratified deposits of the prograding shoreface and beach association. bioturbation in the offshore association varies from weak to intense, but mudstones are commonly completely bioturbated with few remaining physical structures. anconichnus isp., palaeophycus isp., planolites isp. and teichichnus isp. occur in the sandstone beds. interpretation. mudstones with rare laminae of siltstone or sandstone indicate that deposition took place below storm wave base. the thorough bioturbation of the mudstones suggests they were deposited on a shelf with oxic bottom conditions. a higher content of siltstone and sandstone laminae suggests the occasional 318 cm 0 10 20 30 40 50 60 70 80 90 a west lulu-3 3643 m b west lulu-3 3643.9 m c west lulu-3 3645.6 m sb fig. 7. core photographs of selected intervals from the middle paralic wedge of the lulu formation (cal-1b sequence) in west lulu-3 (base lower right, top upper left; for location of core sections, see fig. 18). this core series illustrates the nature of the cal-1b sequence boundary (sb) at 3646 m (c) defined by the erosional base of an estuary channel (facies association 5) cutting into bay/lagoon mudstones (facies association 8). the channel fill sandstones (a–c) show an overall fining-upwards trend and display cross-bedding, flaser lamination and abundant double mud drapes (example arrowed). 319 cm 0 10 20 30 40 50 60 70 80 90 a lulita-1 4504.2 m b amalie-1 5074 m c amalie-1 5110 m d amalie-1 5111 m e amalie-1 5112 m f amalie-1 5118 m sb fig. 8. core photographs of the incised valley-fill of the upper bryne formation (cal-1a sequence). a: erosional surface (sb) marking the base of the incised valley (cal-1a sb) cuts into mottled floodplain mudstones (bat-1b sequence; facies association 3) and is overlain by estuary channel sandstones (facies association 5) showing cross-bedding and mudstone clasts. lulita-1; for location of core section, see fig. 17. b–f: estuary channel sandstones and heterolithic beds in amalie-1 (base lower right, top upper left; for location of core sections, see fig. 17). the largescale cross-bedded sandstones (c, f) with abundant mud drapes and mudstone clasts represent the lower fill of estuary channels (facies association 5). the intervening heterolithic beds (b–e) may represent tidal flats (facies association 8) or fluctuating energy levels in the upper fill of estuary channels (facies association 5). 320 cm 0 10 20 30 40 50 60 70 80 90 a west lulu-3 3647.4 m b west lulu-3 3648.2 m c west lulu-3 3650.8 m d west lulu-3 3651.7 m e west lulu-3 3652.6 m f west lulu-3 3653.5 m g west lulu-3 3654.2 m fig. 9. core photographs of back-barrier deposits of the lower paralic wedge in the lulu formation (cal-1a sequence) in west lulu-3 (base lower right, top upper left; for location of core sections, see fig. 18). the selected core sections illustrate lagoonal mudstones with sandstone interbeds (facies association 8) at the base (g–i), erosively overlain (wavy line, 3654.8 m) by a broadly coarsening-upwards sandstone unit (c–g) – climbing ripple cross-laminated sandstones being succeeded by cross-bedded sandstones with double mud drapes and rare burrows. this coarsening-upwards sandstone unit (facies association 5) shows rootlets (arrowed) towards the top and is capped by a coal bed (c, d); it is succeeded by thoroughly bioturbated sandstones (a, b; mainly teichichnus isp.) representing the upper part of an estuary sand bar that immediately underlies bay/lagoonal deposits spanning the mfs of the cal-1a sequence (not shown in core, see fig. 18). 321 influence of oscillatory currents near storm wave base. sharp-based sandstone laminae are interpreted as storm-sand deposits, and their finer-grained interbeds represent fair-weather sediments and suspension fall-out after storms. deposition took place between storm wave base and fair-weather wave base. facies association 11: prograding shoreface and beach (facies 3–8) description. the prograding shoreface and beach association is represented by up to 12 m thick coarsening-upwards successions of sandstone and subordinate siltstone. the coarsening-upwards successions consist of very fine-grained, hummocky cross-stratified (hcs) and swaley cross-stratified (scs) sandstones with siltstone interbeds in the lower part, overlain by low angle cross-bedded fineto medium-grained sandstones and trough and planar crossbedded fineto coarse-grained sandstones (figs 11b–e, 12). parallellaminated, low-angle cross-bedded and massive fineto coarse-grained sandstones and pebble conglomerates may occur at the top of the successions. the hcsand scs-dominated sandstones occur as sharp-based, laminated beds ranging between a few decimetres and a few metres in thickness, separated by centimetres to decimetres thick siltstone beds. lamination may be gently undulating, and typically intersect and truncate at low angles. individual hummocky crossstratified units may grade into wave-rippled heterolithic siltstone and sandstone. scs sandstones typically occur as thicker amalgamated units that lack the heterolithic sub-units and the silty interbeds. the cross-bedded sandstones occur in poorly defined sets, usually a few decimetres thick. interpretation. the coarsening-upwards successions are interpreted as the deposits of prograding shelf, shoreface and shoreline systems. minor, 2–4 m thick, coarsening-upwards units of typical shoreface deposits may represent wave-influenced mouth bars or ebb tidal deltas. the hcs-dominated units, commonly lowermost in the successions, were deposited by storm wave activity below fair-weather wave base in the offshore transition zone. the scs deposits represent more continuous wave activity on the lower shoreface, whereas the cross-bedded sandstones of the upper part of the succession represent migrating dunes on the upper shoreface. the horizontally laminated and low-angle cross-bedded sandstones uppermost in the successions represent foreshore, beach and strandplain deposits. facies association 12: transgressive shelf and shoreface (facies 9, 10) description. deposits of the transgressive shelf and shoreface association consist of poorly sorted, bioturbated muddy sandstones, sandy siltstones and mudstones and heteroliths, poorly sorted pebbly cm 0 10 20 30 40 50 60 70 80 90 h west lulu-3 3655.7 m i west lulu-3 3656.6 m 322 cm 0 10 20 30 40 50 60 70 80 90 a west lulu-2 3781.6 m b west lulu-2 3782.5 m c west lulu-2 3790 m d west lulu-2 3798.4 m e west lulu-2 3799.2 m sb 323 cm 0 10 20 30 40 50 60 70 80 90 f west lulu-3 3617.4 m g west lulu-3 3618.2 m h west lulu-3 3619.1 m i west lulu-3 3620 m fig. 10. core photographs of the upper paralic wedge in the uppermost lulu formation (cal-1c sequence) illustrating the facies development during the final paralic pulse, prior to regional transgression. a–e: selected core sections from west lulu-2 (base lower right, top upper left; for location of core sections, see fig. 18). lagoonal mudstones and heteroliths of the cal-1b sequence (e) include the cal-1b mfs which correlates distally with marine mudstones of the upper marine wedge (figs 18, 19). the mudstones are abruptly overlain at 3799.5 m (sb, cal-1c sb) by stacked estuary channel fills and bar deposits (b–d; facies association 5). these are capped by sandstones and conglomerates (a, b) deposited in a washover and ravinement complex (facies associations 6, 12) that represents the composite transgressive surface of marine erosion of the cal-1c sequence. f–i: selected core sections from west lulu-3 (base lower right, top upper left; for location of core sections, see fig. 18) illustrating a comparable evolution to that seen in west lulu-2. estuary channel and bar sandstones and heteroliths and lagoonal deposits (g–i) are succeeded by washover/ravinement sandstones and conglomerates (f, g). 324 cm 0 10 20 30 40 50 60 70 80 90 a lulu-1 3594.5 m b lulita-1 4443.5 m c lulita-1 4444.4 m d lulita-1 4445.5 m e lulita-1 4446.3 m f lulita-1 4447.2 m g lulita-1 4448.4 m h lulita-1 4449 m ts ts fig. 11. core photographs of marine shelf and shoreface deposits (facies associations 10, 11) from the lower marine wedge of the lulu formation (cal-1a sequence) in the lulu-1 and lulita-1 wells, close to the basin axis. a: core section from lulu-1 (for location, see fig. 19) illustrating the transgressive surface (ts) overlain by 10–20 cm of thoroughly bioturbated muddy sandstone; this is succeeded by shelf mudstones (facies association 10), including the maximum flooding surface within the interval 3595.2–3594.9 m, grading up into mud-rich heteroliths at the base of a coarsening-upwards prograding shoreface succession (not illustrated here, see fig. 19). b–h: selected core sections from lulita-1 (base lower right, top upper left; for location of core sections, see fig. 19) illustrating a coarsening-upwards prograding shoreface succession (facies association 11). the transgressive surface (ts) at the base (h), overlying lagoonal mudstones, is draped by a thin (1 cm) sandstone, passing abruptly up into structureless shelf mudstones. upwards, the mudstones are interbedded with discrete storm sandstone beds (g) and grade up via heterolithic facies showing hcs (e, f) to sandstones with hcs, scs and cross-bedding (b–d). the maximum flooding surface occurs within the interval 4448–4447.7 m. sandstones, and conglomerates (fig. 10a, f). the deposits are characterised by intense burrowing and a diverse ichnofauna (fig. 11a). interpretation. these sediments were deposited in a shoreface or shallow shelf environment during a transgression. physical structures reflecting the high energy level on the upper shoreface were partly or completely obliterated by burrowing organisms under more tranquil conditions. architecture, depositional environments and sequence stratigraphy a regional unconformity subdivides the bryne formation into two separate parts described here as the lower and the upper bryne formation (figs 3, 4); evidence supporting the recognition of this unconformity is presented below. the architecture and depositional environments of the lower and upper bryne formation and the lulu formation are described here, together with a sequence stratigraphic analysis of these units. lower bryne formation depositional architecture and environments the lower bryne formation consists of floodplain deposits separated by several storeys of channel sandstones. the four most distinct channel units are referred to as units a, b1, b2 and c (figs 3, 4). these channel sandstones can be identified in most wells and probably form laterally continuous sandstone sheets. the lowermost strata of the bryne formation are either finegrained floodplain deposits located below the lowermost channel sandstone (unit a; west lulu-1, west lulu-4) or channel sandstone unit a resting directly and unconformably on triassic or permian deposits (west lulu-2, west lulu-3; fig. 4). in some wells, channel sandstone unit a is a 10–30 m thick multi-storey sandstone section of stacked, finingupwards, 5–15 m thick sandstone units separated by mudstone beds, 1–2 m thick. in other wells, it is a single storey sandstone, 1–2 m thick (fig. 4). the thickness variations may be related to pre-middle jurassic topographic relief. cores are not available from this unit. the two channel sandstone units b1 and b2 are closely associated, usually with the base of b2 lying c. 10 m above the top of b1 (fig. 4). in many wells, the gamma-ray logs of the combined unit b1–b2 show a characteristic fining-upwards – coarsening-upwards – fining-upwards pattern (e.g. amalie-1, 5280–5260 m; west lulu-4, 3768–3740 m; fig. 4). cores are available from unit b2 in the west lulu-3 well (figs 5, 13). channel sandstone unit c is a fining-upwards 10 m thick channel unit recognised in most wells and cored in west lulu-1, west lulu-2 and west lulu-4 (figs 5, 14). unit c consists of cross-bedded sandstone with abundant wood fragments and mud clasts (fig. 5h–k) and an increasing number of clay drapes up-section, some of which are paired. the upper part of the channel unit is heterolithic with decimetre thick sand/mud couplets in west lulu-2 and abundant clay drapes and mud flasers in west lulu-1. the channel units show features that are characteristic of the deposits of sinuous channels. most of the channel bodies have fining-upwards grain-size profiles above erosional bases, they appear to be laterally continuous and regularly spaced mudstone laminae or beds may represent mud drapes on low-angle accretion surfaces (fig. 14). in addition to these features, the cores from sand sheets b2 and c show trough and planar crossbedding, ripple cross-lamination and abundant deformation structures; the basal beds contain intraformational mudstone clasts and wood fragments. the presence of double mud drapes, abundant flaser bedding and decimetre thick sand/mud couplets in channel unit c may indicate that the channel system was influenced by tidal processes. the channel sands were deposited in laterally migrating, sinuous river channels on a coastal plain. the evidence of tidal influence in unit c suggests it may have been connected downstream to an estuary. the upwards increase in tidal influence in this succession may indicate an overall rise in relative sea level. the laterally continuous channel sandstones are separated by up to 50 m thick successions of interbedded mudstone and sandstone (fig. 4). the fining-upwards segments of these successions may appear as an upwards continuation of underlying fining-upwards channel deposits. a mudstone that varies in thickness from a few decimetres to three metres is present at the turnaround point between the fining-upwards and the coarseningupwards segments of the succession (fig. 4). the sandstones form 1–4 m thick units that may fine upwards or show no clear grain-size trends (figs 13–15; fig. 15 faces page 332). the sandstones of both the finingupwards and coarseningupwards segments of the succession show a diverse assemblage of sedimentary structures including climbing ripple lamination, plane 325 326 cm 0 10 20 30 40 50 60 70 80 90 a west lulu-1 3566.9 m b west lulu-1 3567.7 m c west lulu-1 3568.5 m d lulu-1 3575.3 m e lulu-1 3576.2 m f lulita-1 4428.6 m tsme mfs 327 g lulita-1 4431.9 m h lulita-1 4432.9 m i lulita-1 4433.8 m j lulita-1 4434.7 m tsme fig. 12. core photographs of marine shelf and shoreface deposits (facies associations 10–12) from the upper marine wedge of the lulu formation (cal-1b sequence). a–c: sandstones and heteroliths from wave-influenced mouth bar or protected shoreface deposits in west lulu-1 (base lower right, top upper left; for location of core sections, see fig. 18). note the pebble lag (facies association 12) at the transgressive surface of marine erosion (tsme), and the maximum flooding surface (mfs), c. 10 cm higher in the section. d, e: wave-dominated shoreface deposits (facies association 11) in the lulu-1 well showing coarsening-upwards sandstones dominated by scs (base lower right, top upper left; for location of core sections, see fig. 19). f–j: selected cores from a succession of stacked shoreface parasequences (facies associations 10, 11) in lulita-1 (base lower right, top upper left; for location of core sections, see fig. 19). the lower parasequence of heterolithic sandstones (i, j) showing hcs and wave ripple cross-lamination is truncated by a transgressive surface of marine erosion (tsme) and overlain by shelf mudstones (i) that include the maximum flooding surface within the interval 4434.4–4434.1 m. the mudstones grade up into the next prograding shoreface parasequence, comprising heteroliths and sandstones showing hcs and scs (g, h). core section f illustrates the well-sorted swaley cross-stratified sandstones that typically cap the shoreface parasequences. 328 sedimentary/biogenic structures erosional surface parallel bedding/lamination planar cross-bedding trough cross-bedding low-angle cross-bedding hummocky cross-stratification cross-lamination and climbing ripples bimodal current-ripple lamination wave ripples flaser bedding wavy bedding lenticular and silt-streaked bedding mudstone/coal chips disturbed bedding load structures water escape structures synaeresis cracks bioturbation rootlets 5 m ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ amalie-1 3/7-4 lulita-1 lulu-1 west lulu 1 3 2 4 west lulu-3 gr fluvial channel floodplain distal floodplain and lake channel sand b2 baj-1b sb baj-1b mfs cal-1a sb 3711 m 3730 3758 5a 5b 5c 5d 5e 5f ■ ■ ■ ■ ■ ■ siclay sand gr lithology coal claystone siltstone sandstone depositional environments floodplain fluvial channels key surfaces sequence boundary (sb) maximum flooding surface (mfs) søgne basin 10 km fig. 13. sedimentological core log and gamma-ray (gr) log of sequence baj-1b in west lulu-3. the channel sand b2 is located above the baj-1b sb in the basal part of the illustrated section. the baj-1b mfs is placed in the middle of the thick floodplain/lake succession above the channel sand b2. the positions of cores illustrated with photographs here are indicated on the sedimentological logs (e.g. 5a indicates core photograph in fig. 5a). depths of important surfaces, facies changes or core breaks are indicated (in metres below reference level). the accompanying sedimentological legend is also applicable to figs 14–19. lamination and slump structures, as well as dispersed rip-up mud clasts, root traces and pedogenic mottling (fig. 5). the deposits are interpreted as levees, consisting of proximal crevasse splays, small channel fills and small deltas, and more distally as lake, swamp and distal crevasse splay deposits; these sediments are referred to the proximal floodplain association, the lake and distal floodplain association and the vegetated floodplain association. palynological evidence for marine conditions is scarce in these deposits in the søgne basin, although rare marine palynomorphs have been found in the succession separating units b2 and c in west lulu3 indicating that short-lived marine incursions may have occurred. the succession between channel unit c and the cal-1a sb includes both channel sandstones and floodplain deposits (fig. 4). the succession attains a thickness of 30 m in lulita-1, 50 m in west lulu-1, 60 m in amalie-1 and 65 m in 3/7-4. channel sandstones dominate the succession in the west lulu-1 and 3/7-4 wells. cross-bedding is less prominent in the channel sandstones of this succession than in units b2 and c. mudstone clasts and soft-sediment deformation are common, and current ripple and climbing ripple lamination occur in the upper part of channel units (fig. 5). root horizons and pedogenic mottling are abundant in the more fine-grained deposits. fine-grained sediments dominate the succession in the wells closest to the basin axis. in lulita-1, a 14 m thick section of coals and organic-rich mudstones overlie the channel sandstones at the base of the succession (fig. 15). most distinctive among the fine-grained deposits is a 50 m succession of mudstones with thin sandstone and siltstone interbeds in amalie-1. due to the lack of dinoflagellate cysts from this unit, it is interpreted to represent a lacustrine environment. sandy and silty interbeds represent lacustrine delta and lacustrine delta plain/distal floodplain deposits. the 14 m thick coal-bearing section in lulita-1 represents a swamp environment. a relatively deep lake existed in the southern and central part of the basin simultaneously with an active floodplain along the western margin. at least some of the channel sandstones in west lulu-1 and 3/7-4 may represent distributaries of lacustrine deltas (fig. 15). although there is no palynological evidence for marine conditions, short-lived marine incursions of the coastal plain may occasionally have turned the lake into a brackish water lagoon or bay. sequence stratigraphy key surfaces the basinwide extent of both the erosional bases of major channel sandstones and the mudstones that are located at the turnaround points of the fining-upwards – coarsening-upwards successions between the sandstones, suggests that they are not simply the result of autocyclic facies shifts but more likely resulted from regional base-level changes. both channel base diastems and turnaround points can thus be seen as sequence stratigraphic key surfaces. sequence boundaries in the lower bryne formation are defined by channel base diastems of the major, laterally extensive channel sandstones (fig. 4). although key surfaces such as the maximum flooding surface (mfs) and the transgressive surface (ts) do not extend landwards beyond the bay-line (posamentier & vail 1988), non-marine equivalents to the mfs are assumed to occur within widespread lacustrine and floodplain deposits. the presence of an equivalent to the mfs in this setting results from the influence of relative sea-level fluctuations on the groundwater level in the lower coastal plain. surfaces that separate units of amalgamated, laterally extensive channel sandstones from significantly more fine-grained floodplain successions may represent a landwards expression of marine flooding events. in addition, growth of coal-forming peat due to a rise in the groundwater table and associated generation of new accommodation, may be the landwards expression of a marine flooding surface (petersen & andsbjerg 1996). systems tracts in the lower bryne formation, the fluvial sand sheets fine upwards or occur as amalgamated sandstones without a visible grain-size trend (e.g. baj-1a in 3/7-4). the laterally continuous fluvial sand sheets typically found above the sequence boundaries in the alluvial plain deposits represent the lst and the lower part of the tst (figs 3, 4). they are comparable to the low accommodation systems tract of dreyer et al. (1995), and the amalgamated fluvial sand sheet of shanley & mccabe (1991, 1993, 1994) and olsen et al. (1995). channel development was probably initiated during falling or static base level, but the lateral migration of channels during early base-level rise may have caused erosion of a significant proportion of the lowstand deposits. thus, the channel sandstones may largely represent the lower part of the tst. extensive reworking of the floodplain by lateral channel migration during 329 5 m w es t lu lu -1 g r w es t lu lu -2 g r w es t lu lu -4 2. 5 km 1. 8 km g r 36 68m 36 81 38 48 5g 5h 5i 5j 5k m 38 62 38 74 36 92 36 99 36 98 36 85 36 67 36 49m flu vi al c ha nn el s he et ve ge ta te d flo od pl ai n ve ge ta te d flo od pl ai n ba t1b s b ba t1a s b ba t1a m fs c ha nn el s an d c c al -1 a s b c oa l c la ys to ne si lts to ne sa nd st on e fl oo dp la in fl uv ia l c ha nn el s se qu en ce b ou nd ar y (s b) m ax im um fl oo di ng s ur fa ce ( m fs ) li th ol og y d ep os iti on al e nv ir on m en ts k ey s ur fa ce s si c la y sa nd g r si c la y sa nd g r si c la y sa nd g r ■ ■ ■ ■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■■ ■■■■ ■■ a m al ie -1 3/ 74 lu lit a1 lu lu -1 w es t lu lu 13 2 4 ■■ ■■ ■■ sø gn e ba si n 10 k m fi g. 1 4. l o g p an el ( g r a n d c o re l o gs ) o f se q u en ce s b at -1 a a n d b at -1 b i n w es t lu lu -4 , w es t lu lu -2 a n d w es t lu lu -1 . t h e ch an n el s an d st o n e u n it c w h ic h r ep re se n ts a la te ra lly e xt en si ve f lu vi al c h an n el s an d i s lo ca te d a b o ve b at -1 a s b a t th e b as e o f th e se ct io n ; th e b at -1 a m fs r ep re se n ts a b an d o n m en t o f th e ch an n el s ys te m . t h e se ctio n i s d ee p ly i n ci se d b y th e c al -1 a s b , in te rp re te d a s th e b as e o f an i n ci se d v al le y. t h e p o si tio n s o f co re s ill u st ra te d w ith p h o to gr ap h s h er e ar e in d ic at ed o n t h e se d im en to lo gi ca l lo gs ( e. g. 5 k in d ic at es c o re p h o to gr ap h i n f ig . 5k ). d ep th s o f im p o rt an t su rf ac es , fa ci es c h an ge s o r co re b re ak s ar e in d ic at ed ( in m et re s b el o w r ef er en ce le ve l) . fo r fu ll le ge n d , se e fi g. 1 3; i n se t m ap s h o w s th e lo ca tio n o f th e tr an se ct r u n n in g sw –n e . 330 periods of lowstand and early base-level rise may have effectively prevented floodplain aggradation (wright & marriott 1993). the overbank-dominated deposits between the top of the channel sandstone sheets and the mfs, constituting the upper part of the tst, are organised into a fining-upwards succession with a gradually decreasing sand/shale ratio. mudstones of mainly lacustrine origin become increasingly common upwards. soil profiles and root horizons are common (fig. 5g, h, l). the overall fining-upwards trend and accompanying decreasing sandstone/mudstone ratio are interpreted to reflect a sea-level rise that caused a rising watertable and wetter conditions on the coastal plain. the increasing rate of creation of accommodation during rising base level favoured high levels of storage of floodplain sediments and more isolated channel bodies (shanley & mccabe 1993; wright & marriott 1993). the transgressive floodplain deposits are equivalent to the heterolithic unit with isolated fluvial sandbodies of olsen et al. (1995) and the lower part of the high accommodation systems tract of dreyer et al. (1995). the floodplain highstand deposits are separated from the floodplain transgressive deposits by a mfs. the mfs is picked in a mudstone bed that is typically organicrich and can be correlated through most or all the wells in the study area. the highstand floodplain deposits are a coarsening-upwards succession that shows an increase upwards in sandstone/mudstone ratio and sand bed thickness. the lacustrine and distal floodplain mudstones are interbedded with siltstones and sandstones that were deposited as levee deposits, crevasse splays and crevasse deltas. channel sandstones are less common than in the transgressive floodplain deposits. the coarsening-upwards succession of proximal floodplain deposits developed as a result of decreasing rates of base-level rise and accommodation space generation. similar successions in non-marine settings have been referred to the highstand systems tract by shanley & mccabe (1991, 1993), as highstand depositional systems by wright & marriott (1993) and as the uppermost heterolithic interval by olsen et al. (1995). sequences of the lower bryne formation the bryne formation consists of seven sequences of alluvial plain or fluvially-dominated coastal plain deposits. the aalen-1, baj-1 and bat-1 regional sequences of andsbjerg & dybkjær (2003, this volume) are subdivided here into the aalen-1a, aalen-1b, baj-1a, baj-1b, bat-1a and bat-1b sequences. the uppermost part of the bryne formation is included in the cal-1a sequence (see below) that also includes sediments referred to the lulu formation (fig. 4). aalen-1a sequence. the deposits located between the base middle jurassic unconformity and the first intramiddle jurassic sequence boundary (sb aalen-1b) are referred to the aalen-1a sequence. due to onlap of the pre-middle jurassic subcrop the aalen-1a sequence is only seen in wells that penetrate the deepest parts of the middle jurassic. aalen-1b sequence. channel sandstone unit a is referred to the lst/tst of sequence aalen-1b. most of the floodplain deposits separating channel sandstone units a and b1 form the hst of aalen-1b. baj-1a sequence. channel unit b1 and the coarseningupwards unit of floodplain deposits between b1 and b2 form sequence baj-1a. baj-1b sequence. this sequence is made up of channel sandstone unit b2 and the fining-upwards – coarsening-upwards succession of floodplain deposits separating unit b2 from unit c. bat-1a sequence. in most wells, channel sandstone unit c forms the lst/tst of sequence bat-1a. the hst is a unit of coarsening-upwards floodplain deposits. bat-1b sequence. the bat-1b sequence is not present in the westernmost wells where it has been removed by erosion at the cal-1a sb. where present, it consists of channel sandstones above the sb and a finingupwards – coarsening-upwards succession of floodplain and lacustrine deposits (fig. 15). recognition of the intra-bryne regional unconformity in the middle jurassic succession, a number of units and surfaces can be readily correlated across the søgne basin. in most of the bryne formation, sequence boundaries and maximum flooding surfaces are important correlatable key surfaces that are recognisable on both well logs and core logs. the maximum flooding surfaces occur within the mudstone-dominated floodplain deposits between the main channel storeys, being defined by the turnaround point between intervals with increasing-upwards and decreasing-upwards gamma-ray 331 readings (fig. 4). in the lulu formation (see below), the best markers are coal beds, which are easily recognisable on sonic logs and possibly represent ‘initial flooding surfaces’. other surfaces that prove useful for correlation are maximum flooding surfaces (mfs) in the more marine intervals and channel-base diastems in the paralic successions (fig. 4). well-to-well correlation of the coal beds in the lulu formation suggests a sub-parallel arrangement, i.e. that thickness variations are insignificant. these markers are, however, discordant with the marker surfaces in the lower bryne formation (fig. 4). the thickness of the succession between the bat-1a mfs, which is the uppermost key surface of the bryne formation that is easily recognisable in almost all wells of the søgne basin, and the lowermost coal of the lulu formation varies from 12 m in west lulu-4 to 110 m in 3/7-4 and 100 m in amalie-1. this asymmetry may have resulted both from a higher rate of accommodation space generation in the eastern part of the basin due to faulting at the eastern boundary fault and from erosion in the western part of the basin. the boundary between the succession with markers that parallel the lulu formation coal beds and the succession with the non-parallel markers seems to be the cal-1a sb which is a distinct erosion surface at the base of the stacked channel sandstones that dominate much of the upper bryne formation. below this sequence boundary (cal-1a sb), a succession that includes two key surfaces (bat-1b sb and bat-1b mfs) can be recognised in the west lulu-1, 3/7-4, lulita-1, lulu-1 and amalie-1 wells. this succession is missing from the wells in which the stacked channel sandstones above cal-1a sb show their largest thickness (west lulu-2, west lulu-3), suggesting that the missing section is due to erosion at the cal-1a sb and not to up-dip condensation (figs 4, 15). however, in west lulu-4, which is located furthest up-dip of the studied wells, the section between the cal-1a sb and the bat-1a mfs is thin although no significant erosion surface is recognised below the lowermost coal of the lulu formation. this may suggest that the section in this well is condensed rather than missing (figs 4, 16; fig. 16 follows page 332). the basinwide extent of the cal-1a sb erosion surface and the variable but often significant amount of section that seems to have been erosionally removed suggests that the surface is an unconformity with significant relief. the stepwise increase over a relatively short distance of the thickness of the stacked channel sandstones between cal-a sb and the lowermost coal in the lulu formation and the commonly associated increase in the amount of missing section below cal-1a sb suggest that the unconformity represents the basal surface of an incised valley. the base of the lowermost coal in the lulu formation is the first surface above the unconformity that can be correlated to all wells. in west lulu-4, an interfluve surface is inferred within the 10 m thick succession of floodplain deposits that separate the lowermost coal of the lulu formation from the bat-1a mfs. the occurrence of the dinoflagellate cyst impletosphaeridium varispinosum in the basal part of the valley-fill in west lulu-1 and west lulu-3 suggests a late bathonian – early callovian age for the unconformity. this is supported by the presence of the lod of quadraeculina anelliformis in the same interval. upper bryne and lulu formations depositional architecture and environments of the upper bryne formation the upper bryne formation consists of the sediments between the regional unconformity (cal-1a sb) and the first thick coal seam at the base of the lulu formation (figs 4, 16). the regional unconformity is interpreted to form the base of a system of incised valleys in the søgne basin. the southern margin of an e–w-trending valley is inferred to be situated between west lulu-4 and west lulu-2 (fig. 1). the valley axis is interpreted to be close to the wells showing the thickest valley-fills, i.e. west lulu-3 and lulita-1. the position of the northern valley margin is not known due to lack of well data. work on incised valleys from the carboniferous of the north sea area shows that valley-fills with a maximum thickness of 30–40 m as seen in the upper bryne formation, usually correspond to a valley width of at least 5–6 km (hampson et al. 1999). in the wells penetrating the more proximal parts of the valley (west lulu-1, west lulu-2, west lulu-3), massive channel sandstones are present throughout the succession from the basal unconformity to the overlying coal seam (fig. 16). in wells further to the east and to the south, heterolithic and muddy deposits dominate the upper part of the succession. this is most pronounced in amalie-1, which is inferred to penetrate a different valley branch, where 12 m of mudstones and heteroliths are located between the uppermost sandstone body and the coal (fig. 17, following page 332). the channel sands in the proximal wells occur as two to three storeys of amalgamated and stacked channel 332 sandstones separated by fine-grained deposits. each sandstone storey commonly consists of an amalgamated sandstone unit without a distinct grain-size trend although poorly defined fining-upwards trends may occur (e.g. top upper channel storey of west lulu-3; fig. 16). the storeys vary in thickness from approximately 5 m to 18 m (e.g. west lulu-3). the valley-fill is a compound fill (zaitlin et al. 1994), deposited during several minor base-level cycles. it contains significant erosion surfaces in addition to the principal sequence boundary at the base of the incised valley and one or two flooding surfaces on top of the channel storeys. a conglomerate of intraformational mudstone clasts, up to 60 cm thick, may be present at the base of the lowermost channel sandstone (the basal unconformity, cal-1a sb; figs 6l, 16). most sandstones of this succession belong to facies association 5, being characterised by trough and possibly planar cross-bedding, abundant coal and mudstone clasts on foresets and bed boundaries, double mud laminae, flaser bedding and rare bimodal cross-lamination (figs 6c–h, 16). they are interpreted as estuary channel and bar deposits. bioturbated organic-rich sandy mudstones that separate the upper and middle sandstone storeys in west lulu-3 represent lagoonal deposits that developed during a flooding event (figs 6i, 16). most sandstones show some degree of tidal influence, as exemplified by the lowermost and uppermost channel storeys of west lulu-1 (fig. 16). they are interpreted as the fill of tidal channels, deposited mainly as point bars in major estuary channels (see comparable features in fenies & faugères 1998, fig. 8). the middle channel storey of west lulu-3 is an example of a sandstone unit that lacks clear evidence of tidal conditions and may represent a fluvial deposit (fig. 16). the lowermost sandstone bed (1 m thick) in west lulu-2 also shows no evidence of tidal conditions, and may represent a preserved lowstand fluvial deposit (fig. 16). in the 3/7-4 and lulu-1 wells, in the northern and central part of the study area, the sandstone units are thinner, typically 2–5 m thick (fig. 17); they are separated by mudstone-dominated units 2–6 m thick (fig. 17). sandstone units and heteroliths may show a finingupwards pattern, but coarsening-upwards units also occur. sandstones show trough cross-bedding, wavy, flaser and lenticular bedding, bi-directional ripple crosslamination and double mud drapes indicating a tidal environment (fig. 8a). some beds are highly altered by soft sediment deformation and locally by pedogenesis. a few thin coal beds with associated root horizons that occur within the succession suggest periods with vegetation cover. the fining-upwards units represent minor tidal channels. sand-dominated coarsening-upwards units may have been deposited as bay-head or tidal deltas, or as tidal channel bars. heterolithic coarseningupwards beds may represent tidal flat deposits. the combination of an overall tidal setting involving dominantly fine-grained or heterolithic sediments, with only minor channel sandstones as seen in 3/7-4, is indicative of deposition in the outer or marginal part of an estuary (dalrymple et al. 1992). the lulita-1 well displays thicker sand units than in the 3/7-4 well, but thinner and slightly more fine-grained than in west lulu-3; this suggests that lulita-1 was situated close to the channel-dominated axial part of the estuary but downstream from the west lulu wells. in the southernmost part of the basin (amalie-1), the lower 30 m of the succession is dominated by up to 12 m thick sandstone units that fine upwards or show no visible grain-size trend (fig. 17). although locally structureless with only faint trough cross-bedding, these sandstones commonly show trough cross-bedding, with foresets outlined by mud drapes, or ripple cross-lamination and grade up into heterolithic beds with flaser, wavy and lenticular bedding (fig. 8c–f). mudstone flakes are abundant in some sandstone beds; waterescape structures are also common in places. the uppermost 11 m of the succession mainly consist of finegrained heteroliths and mudstone (fig. 8b). amalie-1 is located approximately 12 km south of the inferred valley axis trending from west lulu-3 to lulita-1 so that amalie-1 is thought to penetrate the valley-fill deposits of a separate, n–s-trending valley. the occurrence of fine-grained heterolithic deposits in the upper part of the succession in the wells to the east and south-east may indicate an up-dip shift of facies due to a relative sea-level rise, with tidal flat and lagoonal facies becoming dominant in the lower reaches of the valley. however, this may also have resulted from an autocyclic shift of facies in an outer estuary environment, where widespread tidal flats bordered tidal channels. it is also possible that the shift to fine-grained sedimentation represents a change to deposition unconfined by valley walls when infill was complete in the lower reaches of the incised valleys. depositional architecture and environments of the lulu formation viewed in an east–west transect across the western and central parts of the søgne basin, the lulu formation 333 consists of three eastwards-thinning wedges of mainly paralic deposits and two westwards-thinning wedges of shallow marine and coastal deposits (figs 3, 18, 19, 20b; figs 18, 19 follow page 332). at the top of the lulu formation is a transgressive unit of shallow marine and coastal deposits, a few metres thick. two regionally extensive coal seams/coal zones can be traced across the basin; one separates the lulu formation from the bryne formation and the other divides the lulu formation into a lower and an upper part of almost equal thickness (fig. 20b). coals the basal coal seam, locally split into several thinner seams, is up to 5 m thick in the søgne basin. petersen & andsbjerg (1996) have described this basal coal from west lulu-2 as seams r1 and t2, which record a relatively dry peat-forming environment below (r1) succeeded by a waterlogged peat-forming environment above (t2). the coal seam is a single, almost structureless coal bed in west lulu-1 and 3/7-4 but is represented by two or three distinct coal beds in west lulu-2, west lulu-3 and west lulu-4 and by a zone of interbedded thin coals and lagoon and marsh sediments in lulu-1 and amalie-1. the upper coal seam divides the lulu formation into a lower and an upper part of almost equal thickness (figs 18, 19). this seam is located in the middle wedge of the three eastwards-thinning paralic wedges and can be correlated throughout the study area. it occurs as a single 0.2–0.4 m thick coal bed in the central and southern part of the søgne basin, and as two coal beds (max. 2 m thick) in the westernmost wells. paralic wedges the three paralic wedges thin from the west towards the east. excluding the lower coal zone, the lower paralic wedge is 11–13 m thick in the west lulu-1, west lulu-2 and west lulu-3 wells (fig. 18), 7 m thick in 3/7-4 and 1–3 m thick in the lulita-1, lulu-1 and amalie-1 wells (fig. 19). it is bounded below by the top of the lower coal seam/coal zone and above by a distinct flooding surface that separates it from deposits of the lower marine wedge. in the western part of the basin, the geometry of the middle paralic wedge is poorly constrained. the marine flooding surface that separates the middle paralic wedge from the upper marine wedge can be placed at two alternative positions in the west lulu-3 well – at a gammaray pick in an uncored section at 3639 m and at a waveinfluenced heterolithic sandstone bed at 3626 m (fig. 18). the latter interpretation implies an even more dramatic westwards-thickening of the middle paralic wedge than that seen for the lower paralic wedge. the former is preferred here, i.e. the flooding surface is placed at approximately the same level, above the upper regional coal marker, as in west lulu-1 (fig. 18). in west lulu-2 and west lulu-3, the lower boundary of the wedge is a channel-base diastem (cal-1b sb). in all other wells, the boundary is picked at the top of the beach deposits that terminate the coarsening-upwards marine succession of the lower marine wedge. the thickness of the wedge in the westernmost wells is 3 m in west lulu-1, 6 m in west lulu-2 and 10 m in west lulu-3. in 3/7-4, lulita-1, lulu-1 and amalie-1, the wedge has a constant thickness of approximately 2 m. the upper paralic wedge attains a thickness of 20 m in west lulu-3, 21 m in west lulu-2, 8 m in 3/7-4, and 3–6 m in lulita-1, lulu-1 and amalie-1. much of the wedge is assumed to have been removed by faulting in west lulu-1. the lower boundary is a distinct channel-base diastem (cal-1c sb) in the west lulu-2 well but is located at the shift from the shoreface and foreshore deposits of the upper marine wedge to the overlying strandplain and back-barrier deposits in the wells further to the east. in west lulu-3 and west lulu-1, the boundary has been placed at the erosional base of a coarsening-upwards sandstone unit interpreted as tidal bar or mouth bar deposits of a prograding bay-head delta, sitting below the cal-1c sb. the upper boundary is the final marine flooding surface below the transition to the offshore mudstones of the lola formation. this surface is placed at the base of a 1.5 m thick transgressive sandstone bed in 3/7-4, and in west lulu-2 and west lulu-3 at the base of a transgressive conglomerate/pebbly sandstone unit. in the west lulu-2 well, the paralic wedges are dominated by up to 10 m thick storeys of stacked sandstone units that fine upwards or show no grain-size trends. in the west lulu-1 and west lulu-3 wells, the paralic wedges are characterised either by coarsening-upwards mudstone–sandstone successions or by sandstones showing no clear overall grain-size trends. both coarsening-upwards and fining-upwards sandstones occur in 3/7-4. the sandstones show trough cross-bedding, current ripple cross-lamination and double mud drapes. climbing ripple cross-lamination and bioturbation (abundant teichichnus isp. burrows, less common planolites isp. and skolithos isp.; fig. 9a, b) occur frequently in the coarsening-upwards units and the units showing no 334 grain-size trends. flaser bedding, soft-sediment deformation structures and beds with abundant coal and mudstone clasts are characteristic of the fining-upwards sandstone units. heterolithic intervals show wavy and lenticular bedding. the coarsening-upwards successions are interpreted as tidal bar deposits or mouth bar deposits of prograding bay-head deltas. the sandstone units that fine upwards or show no grain-size trends represent the fill of major estuary channels. minor finingupwards sandstone beds within, and typically near the top of, coarsening-upwards successions, may represent bay-head delta distributary channels. in both 3/7-4 and lulita-1, the c. 1 m thick sandstonedominated lower part of the upper paralic wedge consists of a succession of parallel-laminated and low-angle cross-bedded sandstones interpreted as strandplain deposits. this is overlain by a heterolithic unit, up to 2.5 m thick, showing flaser, wavy, and lenticular bedding deposited in a low-energy estuary or lagoon environment. in 3/7-4, the heterolithic succession is abruptly overlain by a 3 m thick unit of stacked sandstones with coal and mud clasts, which fines upwards into an organic-rich, heterolithic mudstone with abundant roots, and finally a coal bed. these sandstones are interpreted as the fill of a minor distributary channel. in the lulu-1 and amalie-1 wells, located in the central and southern part of the basin, the lower paralic wedge is represented solely by coals with interbedded clastic sediments of the lower coal zone. paralic deposits above the coals have been reworked and incorporated in the lower marine wedge during transgression. the middle and upper paralic wedges consist of 1–1.5 m of poorly sorted, structureless sandstone with abundant coal debris and root traces; the sandstone may show irregular ripple cross-lamination. these deposits are interpreted as mainly strandplain deposits. the upper coal seam/coal zone is situated within the middle paralic wedge. the occurrence of minor coal beds and root horizons indicates periods with vegetation cover and peat accumulation. marine depositional wedges the two marine wedges both thin towards the west. the lower marine wedge attains a thickness of 7.5–11.5 m in the lulita-1, lulu-1 and amalie-1 wells, c. 3 m in 3/7-4 and west lulu-1, and only about 1 m (preserved thickness) in west lulu-2 and west lulu-3 (figs 18, 19). the upper marine wedge is approximately 10 m thick in the lulita-1, lulu-1 and amalie-1 wells and attains a thickness of 6 m in 3/7-4 and 5 m in west lulu-1. in west lulu-3 and west lulu-2, it is represented by an approximately 1 m thick mudstone bed. in both lulu-1 and amalie-1, the basal part of the lower marine wedge consists of an erosionally based, fining-upwards silty sandstone unit, 40–60 cm thick. the sandstone becomes increasingly heterolithic or muddy upwards, but primary structures have been obliterated by pervasive bioturbation (fig. 11a). the sandstone is interpreted as a transgressive shelf deposit. it is overlain by a unit of shelf mudstones that varies in thickness from about 0.5 m in lulu-1 to approximately 11 m in amalie-1, where it shows a coarsening-upwards trend. above the mudstone unit in lulu-1, lulita-1 and amalie-1 is an 8–12 m thick, coarsening-upwards succession. the basal part of the coarsening-upwards succession consists of mudstone–sandstone heteroliths, showing lenticular and parallel bedding/lamination and wave ripple cross-lamination in sand laminae. it is overlain by sand-dominated heteroliths with abundant hummocky cross-stratification, sandstones with low-angle and swaley cross-stratification and trough cross-bedding, and uppermost by parallel bedded and low-angle crossbedded sandstone (figs 11a–f, 19). this represents a progradational succession from offshore transition to lower shoreface sediments overlain by upper shoreface and beach deposits. the wave-dominated sediments of the lower marine wedge can be traced to the west in west lulu-1 and 3/7-4 as a unit up to 4.5 m thick. in west lulu-1, a pebble veneer interpreted as a wave ravinement lag defines the base of the wedge. this is succeeded by a 0.5 m thick unit of hummocky cross-stratified and wave-rippled sandstone overlain by a thin mudstone and a 3 m thick coarsening-upwards sandstone unit dominated by swaley cross-stratification and low-angle cross-bedding with abundant teichichnus isp. burrows. in the 3/7-4 well, the base of the wedge is picked at a flooding surface below which the uppermost paralic deposits are strongly bioturbated. the deposits of the marine wedge form several 0.5–2.5 m thick, coarsening-upwards sandstone units showing wave ripple lamination, lowangle cross-bedding and hummocky and swaley crossstratification. teichichnus isp. burrows are common. in both west lulu-1 and 3/7-4, this succession is interpreted as a condensed shoreface or a wave-influenced mouth bar. in west lulu-2 and west lulu-3, the westernmost correlative of the lower marine wedge consists of a few metres of mudstone. thorough bioturbation with abundant teichichnus isp. burrows in the top of the underlying paralic deposits indicates that a marine or brackish flooding event preceded deposition of the 335 mudstone. erosion at the cal-1b sb has removed all shallow marine or coastal deposits of the lower marine wedge that may have overlain the mudstone. in the upper marine wedge, the shoreface succession in lulita-1 and lulu-1 is represented by three stacked, 1.2–4 m thick, coarsening-upwards parasequences separated from each other by minor flooding surfaces. further to the north in 3/7-4, the 10 m thick upper marine wedge has a similar architecture with three parasequences (2 m, 2 m, 6 m) separated by distinct flooding surfaces. each parasequence consists of a basal unit of heterolithic sandstones and mudstones that coarsens upwards to sandstones dominated by wave-generated structures, suggesting a shoreface origin (fig. 12d–j). in 3/7-4, the facies assemblage indicates deposition in a mixed waveand tide-dominated environment, possibly a tidally influenced mouth-bar or an ebb tidal delta. in west lulu-1, the upper marine wedge consists of a 3.5 m thick coarsening-upwards unit (fig. 12a–c). it comprises heterolithic siltstone and sandstone showing parallel lamination, wave ripple lamination, lenticular bedding, and hummocky cross-stratification, and sandstone showing low-angle cross-bedding and possibly swaley cross-stratification. both in west lulu-1 and 3/7-4, teichichnus isp. and diplocraterion isp. burrows are common. the abrupt upwards termination of the wedge in west lulu-1 may be caused by a normal fault. the succession in west lulu-1 represents a wavedominated environment, interpreted as a progradational shoreface deposited in an area with limited accommodation, or a wave-influenced mouth bar. final transgressive deposits the channel and estuarine bar sandstones of the upper paralic wedge are erosionally overlain by an up to 4 m thick unit consisting of sandstones and pebble conglomerates. the base of this unit is commonly outlined by a pebble veneer draping an erosion surface. in west lulu-3, this unit includes several erosionally-based beds, up to 10 cm thick, of graded clast-supported pebble conglomerate (fig. 10f). interbedded with the conglomerates are beds of well-sorted sandstone and pebbly and granule-rich sandstone. in 3/7-4, this part of the succession is represented by a 1.5 m thick heterolithic sandstone dominated by wave ripple lamination and wavyand lenticular bedding. the erosional surface that lies at the base of the conglomerates in west lulu-2 and west lulu-3 and separates the tidally influenced sandstones from the overlying fine-grained marine sediments in lulu-1 and 3/7-4, is interpreted as a transgressive surface of marine erosion (tsme) or ravinement surface. the coarse-grained sediments above the ravinement surface in west lulu-2 and -3 were deposited as beach and shoreface deposits during transgression (bourgeois & leithold 1984). sediments of that grain size are rare in the underlying succession, and they are therefore interpreted as the result of storm-wave reworking of coarse fluvial sediments supplied to the near-shore zone. the graded pebbly sandstones sandwiched between the conglomerates and the overlying marine mudstones in west lulu-2 (fig. 10a, b) represent rapid deposition of sediment eroded by waves breaking on the shoreface (bourgeois & leithold 1984). sequence stratigraphy of the upper bryne formation and the lulu formation key surfaces within the mainly estuarine deposits of the upper bryne formation, flooding surfaces (fs) separate stacked channel sandstones from overlying lagoonal or marine mudstones. channel-base diastems that can be correlated throughout the incised valleys possibly represent sequence boundaries of higher order sequences although no attempt has been made to subdivide that part of the succession further. in tidally dominated paralic units in the lulu formation, sequence boundaries are defined by channel-base diastems (fig. 18). in marine intervals, the sequence boundaries occur as indistinct surfaces that separate beach deposits from overlying strandplain deposits (fig. 19). in the wells located in the central and southern parts of the study area, the basal sequence boundaries of the cal-1b and the cal-1c sequences are placed immediately above the beach deposits that form the top of the prograding shoreface successions (fig. 19). the shift from beach deposits to the overlying laterally extensive coalbearing or rooted beds indicates a basinwards shift of facies. in 3/7-4, the cal-1b sb (3460 m) is identified at the base of a rooted sandstone bed sitting on top of the condensed shoreface or mouth bar succession that comprises the hst of cal-1a; in this well, the cal-1c sb is placed at the base of a rooted channel sandstone (3449 m). in west lulu-1, the cal-1b sb is represented by a bed of pebbly sandstone (core rubble) at 3572 m. in the west lulu-2 and west lulu-3 wells, the lulu formation is dominated by stacked channel sandstones; the sequence boundaries of the cal-1b and cal-1c sequences are placed at the base of coarse-grained 336 channel sandstones in these wells. the cal-1b sb cuts into lagoonal mudstones in west lulu-3 (fig. 7c) and the cal-1c sb cuts into lagoonal mudstones in west lulu-2 (fig. 10e; 3799 m). these lagoonal mudstones may include the cal-1a and cal-1b maximum flooding surfaces although these surfaces may have been truncated by the channel base diastem. the cal-1b sb is located at the base of a thin sandstone at 3808 m in west lulu-2; the cal-1c sb is picked at the erosional base of a channel sandstone unit in west lulu-3 at 3625 m. a marine flooding surface subdivides this sequence (cal-1c) into a lower unit dominated by paralic sandstones (uppermost lulu formation) and a mudstonedominated upper unit (lowermost lola formation). in some wells, a transgressive surface of marine erosion (ravinement surface) can be seen immediately below the flooding surface. an erosion surface that separates shelf mudstones and shoreface transition heteroliths from overlying shoreface sandstones is interpreted as a regressive surface of marine erosion (rsme); it is located at 3582 metres in lulu-1, 4433 metres in lulita-1, 3457 metres in 3/7-4 metres and at 3565 metres in west lulu-1 (figs 18, 19). the shoreface sandstones above the rsme are referred to the falling stage systems tract (fsst). systems tracts in the valley-fill deposits that constitute the upper bryne formation, lst-deposits of the cal-1a sequence, if present, are to be found among the massive channel sandstones that dominate the valley-fill. however, most of these channel sandstones show clear evidence for tidal processes, and are referred to the tst, recording an increase in the rate of relative sea-level rise. during the lowest sea-level stand, incised valleys acted as conduits for sediment by-pass, and much of the fluvial sediment deposited within the valley may have been eroded and shed further basinwards. a preserved fluvial sandstone bed, 1 m thick, at the base of the valley-fill succession in west lulu-2 may represent the lst of the cal-1a sequence (fig. 16). in the cal-1b and cal-1c sequences, channel deposits directly overlying the sequence boundaries in the west lulu-2 and west lulu-3 wells show evidence of strong tidal influence (figs 7a, b, c, 10d). the thin sedimentary section between the cal-1b sb and the overlying coal in the wells further east does not show any diagnostic sedimentary structures. a typical tst in the upper bryne and lulu formations in the west lulu area consists of a lower succession dominated by tidally influenced fluvial channel and estuary channel sandstones, some of which may have been deposited in an incised valley, and an upper succession of outer estuary and lagoonal deposits. in the wells closer to the basin axis, estuary channel deposits are only important constituents of the tst when located in an incised valley. otherwise, the tst in this area is dominated by outer estuary, marine bay, and transgressive shoreface and shelf deposits; a ravinement surface or transgressive surface of marine erosion (tsme) normally separates the lower estuarine part of the tst from transgressive shoreface deposits. the uppermost succession of the tst normally wedges out in a basinwards direction. the tst is bounded above by the maximum flooding surface (mfs) represented by shelf or lagoonal mudstones. in the marine successions, the hst is a coarseningupwards succession of shelf, shoreface and beach deposits. the hst wedges out in a landwards direction where the succession consists of bay-head and tidal delta deposits overlying lagoonal or bay mudstones reflecting a rapid, progradational infilling of estuaries or bays. the hst is truncated above by a sequence boundary or in some cases by a regressive surface of marine erosion (rsme). truncation at the cal-1b sb causes the cal-1a hst to be absent from west lulu-2 and west lulu-3. similarly, the cal-1b hst is missing in west lulu-2 due to erosion at the cal-1c sb (fig. 18). a significant erosional break within the coarseningupwards succession of regressive shoreface deposits in the cal-1b sequence in lulu-1 and lulita-1 suggests that the upper shoreface and foreshore deposits above the break were deposited during a fall in relative sea level, which caused wave erosion of the already deposited lower shoreface and shelf sediments (plint 1988). the deposits between the erosional break (rsme) and the next sequence boundary are referred to the falling stage systems tract (fsst). the fsst consists of coarsening-upwards shoreface, estuary mouth, foreshore and beach deposits; their formation and preservation was dependent on the balance between sea-level change and subsidence. in addition to their occurrence in lulu-1 and lulita-1, regressive shoreface deposits may possibly be referred to a cal-1b fsst in the west lulu-1 and 3/7-4 wells (figs 3, 19). sequences of the upper bryne and lulu formations the three sequences cal-1a, cal-1b and cal-1c cover the uppermost part of the bryne formation, the lulu formation and the lowermost part of the upper jurassic lola formation (fig. 20). 337 338 cal-1c sb cal-1b sb cal-1a sb cal-1c cal-1b cal-1a estuary channels estuary channels transgressive shoreface mouth bar deposits incised valley fill prograding shoreface a b cal-1a sb bat-1b sb bat-1a sb baj-1b mfs baj-1b sb aalen-1b sb aalen-1a sb floodplain floodplain floodplain channel sand c channel sand b2 lacustrine deposits channel sand b1 channel sand a west east depositional environments floodplain fluvial channels lacustrine mires, swamps lagoon and tidal flats estuary and tidal channels, bay-head deltas shoreface, mouth bars and washovers marine shelf key surfaces sequence boundary (sb) maximum flooding surface (mfs) fig. 20. schematic representation of the sequence stratigraphy and stacking patterns of the lower bryne formation (a) and the upper bryne formation and lulu formation (b). possible lst deposits and the lower tst of cal-1a are represented by incised valley-fill deposits, referred to the upper bryne formation (figs 16, 17). the remainder of the cal-1a sequence is made up of the lower paralic wedge and the main part of the lower marine wedge; these are assigned to the upper part of the tst and the hst (figs 18, 19). in the cal-1b sequence, possible lst deposits and most of the tst are represented by the middle paralic wedge; the upper marine wedge is assigned to the uppermost part of the tst, the hst and the fsst. the upper paralic wedge is referred to the lst(?) and the lower part of the tst of cal-1c. the remainder of the cal-1c sequence occurs within the lowermost lola formation, where the hst is represented by a progradational unit of shallow marine, strongly bioturbated mudstone and silty sandstone. depositional history and palaeogeography aalenian(?) – late bathonian between the aalenian/early bajocian and the middle/ late bathonian, when base level was low, the study area was dominated by an alluvial plain with laterally migrating, sinuous rivers that swept most of the floodplain (fig. 21a). the presence of stacked, amalgamated channel sandstones in the west lulu-1 and 3/7-4 wells in the vicinity of the lulu salt structure and its northwards extension suggest that this area, in particular, was favoured by channels. deposition took place on a coastal plain, where the upstream effects of tidal processes were occasionally felt in the river channels. recurrent periods of rising base level resulted in the abandonment of the large river channels. the area changed into a wet floodplain environment dominated by ponds and minor channels. at the time of maximum flooding, extensive lakes occupied the axial part of the basin and other topographic lows (fig. 21b). brackish or fully marine waters may have entered the basin on occasion to form shallow bays or lagoons, particularly in the southern part of the danish central graben. regional drainage was from the north to the south where marine conditions existed in the dutch part of the central graben until the early bathonian (van adrichem boogaert & kouwe 1993; hengreen et al. 2003, this volume). during periods with a decreasing rate of base-level rise, lacustrine deltas and crevasse splays filled in the lakes and lagoons, and a depositional environment dominated by laterally migrating rivers was re-established. during the bathonian, more perennial lakes may have existed in the southern part of the søgne basin, while swamps developed in the northern and the central part of the basin. late bathonian – callovian during formation of the base cal-1a sb, major incised valleys were cut both at the western fringe of the søgne basin and in the south-eastern part of the basin close to the basin axis (amalie-1; fig. 21c). late bathonian – earliest callovian datings have been obtained from the lower part of the incised valley-fill. broad estuaries developed in the lower reaches of the incised valleys during relative sea-level rise. deposition took place mainly in major channels in the more proximal parts of the valleys, and in outer estuary environments characterised by tidal flats, minor tidal channels and flood tidal deltas closer to the basin centre (fig. 21c). locally, a final phase of valley-fill is evident, characterised by fine-grained sediments deposited in tidal flat and lagoonal environments (e.g. amalie-1; figs 17, 22b). the general increase upwards in tidal influence and preserved thickness of channel storeys seen in many valley-fill deposits suggests that deposition took place during rising sea level. once the incised valleys were completely filled, sedimentation was no longer laterally confined (fig. 22). the estuary environment was replaced by a low-energy lower coastal plain, which was dominated by extensive coal-forming mires and swamps that extended over both the infilled valleys and the former interfluve areas. mire aggradation resulted in thick coal-generating peat deposits in the western part of the søgne basin, whereas coastal swamps caused the formation of thin coals and coaly mudstones in the central and southern parts of the basin. the resulting coal seam records a stepwise increase in marine influence with time, as a continuously waterlogged environment with occasional seawater incursions succeeded a relatively dry peat-forming environment (petersen & andsbjerg 1996). the growth of extensive peat-forming mires and swamps ended as a result of the combined effects of continued sea-level rise, causing transgression in the north-east, and clastic influx from up-dip sources in the west. deposition of shelf mud began in the central part of the søgne basin following transgression, while a lagoonal/estuarine environment was established in the western part of the basin (fig. 23a). continued 339 340 amalie-1 w. lulu-3 w. lulu-1 lulu-1 lulita-1 lulita-1 w. lulu-4 3/7-4 w. lulu-2 amalie-1 w. lulu-3 w. lulu-1 lulu-1 w. lulu-4 3/7-4 w. lulu-2 major lake lacustrine deltas10 km 10 km 10 km alluvial plain with river and ox-bow lake floodplain and lacustrine delta lake n n n 3/7-4 w. lulu-3 w. lulu-1 w. lulu-2 lulu-1 amalie-1 w. lulu-4 estuary interfluvial coastal plain terrace estuary with marsh and sandbar coastal plain with river lulita-1 a c b incised valley tidally influenced river incised valley fig. 21. a: palaeogeographic map for the lower bryne formation (aalenian – late bathonian). when regional base level was low, laterally migrating sinuous rivers dominated the floodplain. the absence of the lowermost bryne formation in the lulu-1 well is attributed to uplift related to the underlying salt dome, forming a weak positive feature in the aalenian – early bathonian. b: palaeogeographic map for the upper levels of the lower bryne formation (bathonian), depicting the high base-level scenario involving distal floodplain, lake and lacustrine delta depositional systems. the lake may have been influenced by marine incursions with the development of brackish bay/lagoonal conditions. note that this palaeogeographic scenario is also applicable, in general, to times of high base level in the aalenian – early bathonian although the detailed distribution of environments will have been modified by the lulu-1 positive feature. c: palaeogeographic map for the upper bryne formation (late bathonian – earliest callovian). major incised valleys with estuary channels and tidally influenced river channels representing the lst and lowermost tst of the cal-1a sequence. 341 interfluve estuary in incised valley estuary channel bars early valley fill – lst/lower tst interfluve estuary in incised valley lagoon marsh a drowned estuary late valley fill – lower tstb lower bryne fm lower bryne fm fig. 22. block diagram showing the inferred palaeogeography during deposition of the uppermost bryne formation. deep incision, creating the cal-1a sequence boundary, resulted in two incised valley systems, the confluence of which is depicted here. the w–e transverse system, draining the hangingwall slope, is encountered particularly in the west lulu wells whereas the rift-axial system, trending s–n parallel with the main boundary fault, is represented by the amalie-1 section (fig. 21c). in their lower reaches, as depicted here, the valleys were estuarine in nature and were progressively drowned; this evolutionary phase is recorded in the lower tst of the cal-1a sequence. transgression towards the west and south caused reworking of coastal and back-barrier deposits, while a progressively thicker succession of back-barrier deposits was preserved below the transgressive surface of marine erosion. the overall transgression of the søgne basin was interrupted on at least two occasions by regressive phases caused by periods of relative sea-level fall or stillstand. each regressive phase began with prograding bay-head deltas infilling lagoons and estuaries in the western parts of the area. when infilling of lagoons and estuaries was complete, sediment began to bypass the coastal zone and was supplied to the shoreface. shoreface sediments prograded into the deeper parts of the basin forming a wedge of shallow marine and coastal deposits (fig. 23b). if the regressive phase was associated with a sea-level fall, the decreasing accommodation caused increased wave scour on the inner shelf, and rapid progradation of the shoreface. a thin sheet of strandplain sediments deposited behind the prograding coastline is indicative of the completion of infilling. thin extensive coal deposits that overlie the strandplain deposits indicate a shift from regression to renewed transgression. during the final transgression of the area, probably in the late callovian, the top of the coastal plain deposits was eroded by wave action resulting in the formation of a ravinement surface. the rapid transition from paralic sediments to offshore mudstones and siltstones indicates a rapid transgression across a low-gradient 342 shelf and shoreface barrier bar and tidal inlet amalie-1 w. lulu-3 w. lulu-1 lulu-1 lulita-1 lulita-1 w. lulu-4 3/7-4 w. lulu-2 coastal plain 3/7-4 w. lulu-4 amalie-1 lulu-1 tidal inlet 10 km 10 km beach and shoreface strand plain and delta shoreface transition and shelf rivermarsh and lagoon coastal plain with river and bay-head delta n n w. lulu-3 w. lulu-1 w. lulu-2 a b fig. 23. a: palaeogeographic map for the lulu formation (callovian) depicting transgressive shoreface, barrier coast and coastal plain settings, a scenario recorded by the upper tst of sequences cal-1a and cal-1b. b: palaeogeographic map for the lulu formation (callovian). this scenario, involving progradational shoreface, beach ridge plain and alluvial plain settings, is inferred from the hst/frst of sequences cal-1a and cal-1b. coastal plain. sediment sources were effectively removed from the vicinity of the study area. deposition of paralic and shallow marine sandstones of the bryne formation was terminated when the basin entered the rift climax phase. sediment supply was no longer sufficient to keep pace with the increased rate of subsidence, and deposition of the lola formation shelf mudstones took over. discussion the middle jurassic deposits of the danish central graben form part of a major system of alluvial plain, coastal plain, delta plain and shallow marine deposits that extends over large tracts of the north sea area. during the earliest middle jurassic, large-scale regional uplift, the ‘north sea doming event’ and the subsequent dome collapse affected a large part of this area (whiteman et al. 1975; eynon 1981; ziegler 1990; underhill & partington 1993). the pre-middle jurassic deposits in the danish central graben are cut by a major unconformity that separates the middle jurassic succession from the lower jurassic fjerritslev formation in the southern and central part of the danish central graben and from triassic and permian rocks in the søgne basin (andsbjerg et al. 2001). in contrast to the middle jurassic in much of the north sea area, fully marine deposits have not been found in the pre-callovian of the søgne basin. in the northern viking graben, the lower part of the brent group is dominantly marine. in both the cleveland basin of eastern england and in the norwegian–danish basin, marine deposits are well-represented in the lower part of the middle jurassic succession. in the danish central graben, thin mudstones that yield dinoflagellate cysts occur in the lower part of the middle jurassic succession more than 50 km south of the søgne basin (andsbjerg 1997). further south in the dutch part of the central graben, the aalenian – lower bathonian succession comprises marine mudstones of the werkendam formation (van adrichem boogaert & kouwe 1993; herngreen et al. 2003, this volume). in the northern part of the central graben (gatliff et al. 1994) and in the moray firth (maclennan & trewin 1989), middle jurassic deposits older than the bathonian or latest bajocian seem to be absent, probably due to their location near an early middle jurassic uplift centre. the occurrence of marine lower middle jurassic deposits in the southern part of the central graben simultaneously with non-marine deposits in the søgne basin and the possible absence of lowermost middle jurassic rocks in the northern central graben suggest that regional drainage patterns within the central graben was from the north towards the south, being strongly influenced by uplift patterns. in the late middle jurassic, the appearance of marine deposits in the søgne basin simultaneously with nonmarine deposition to the south and south-west suggests a significant change in regional slope and drainage patterns. during the callovian, drainage in the danish central graben was from the west and south-west, down newly developed hangingwall slopes, and possibly from uplifted areas in the southernmost part of the north sea. most of the middle jurassic succession was deposited during the early stages of rift-related subsidence in the søgne basin. at some stratigraphic levels, sediments and facies patterns show an asymmetric distribution across the basin. between the fluvial sand sheets, lacustrine and distal floodplain deposits tend to dominate in the wells of the central and southern part of the basin closest to the main boundary fault whereas proximal floodplain deposits are dominant in the western part of the basin (fig. 20b). the available well data do not suggest a preferred positioning of fluvial channels close to the main boundary fault, but a tendency to amalgamation and thickening of channel sands is seen in the wells near the lulu salt structure and its northwards extension. in a relatively arid environment, fluvial channel sands parallel to the basin axis would show a tendency to cluster near the main boundary fault (alexander & leeder 1987; leeder & gawthorpe 1987). that this is not the case in the søgne basin may be explained by a setting on a coastal plain with a high groundwater level resulting in the development of lakes and wet floodplain environments in the deep parts of the basin as a response to subsidence at the main boundary fault. under such conditions, during periods of active subsidence, transverse fluvial systems would be located on the hangingwall slope draining into the axial lakes (alexander & leeder 1987; leeder & gawthorpe 1987). only during periods of tectonic quiescence could large, longitudinal fluvial systems develop after lake-infilling was complete. in the paralic to shallow marine succession in the upper part of the middle jurassic section, depositional units show a spatial partitioning such that paralic sediments dominate towards the west, deposited mainly during rising sea level, and offshore–shoreface sediments dominate towards the east, deposited during highstand and possibly early fall in sea level. such landwards partitioning of paralic deposits during trans343 gression and seawards partitioning during highstands has been described previously by ravnås & steel (1998) and is analogous to the ‘reciprocal’ style of sedimentation described from the gallup sandstone of new mexico by nummedal & molenaar (1995). coal beds that were deposited on a low-gradient coastal plain, are overlain by paralic deposits of the tst that show a progressive increase in thickness towards the west or up-dip on the hangingwall slope (fig. 20b). in contrast, coastal and shallow marine deposits of the hst and fsst overlying the paralic paralic wedge thicken towards the east or down-dip on the hangingwall slope. the tsme, or in some places the mfs, that separates the two wedges, thus shows a significantly higher gradient than the coal below the paralic wedge. the widespread thick coals at the base of the succession indicate initial conditions characterised by a low gradient and negligible sediment input. a transgression in that setting would be expected to be a rapid, low-angle non-accretionary transgression (helland-hansen & martinsen 1996). however, the angular difference between the coals (the original depositional surface) and the transgressive shoreline trajectory represented by the tsme and the mfs above the preserved wedge of paralic transgressive deposits is suggestive of a change in slope before or during the early phases of transgression. thus an accretionary transgression took place, possibly after an initial phase of low-angle non-accretionary transgression (helland-hansen & martinsen 1996; fig. 20b). the westwards-thickening wedge of paralic sediments that formed during transgression left eastwards-increasing accommodation space unfilled at the time of maximum flooding. this accommodation space was filled by prograding mainly shallow marine deposits during the subsequent highstand or possibly the falling stage, resulting in the reciprocal distribution pattern of westwards-thickening tst deposits and eastwards-thickening hst/fsst deposits (fig. 20b). the nature of the sediment partitioning and the occurrence of an aggradational transgression rather than a low-angle non-aggradational transgression on the hangingwall slope, as may be expected from the initial conditions, can be explained by tectonic influence. faultinduced tilting of the original depositional surface would have caused a slower transgression of a steeper slope and a concentration of the available volume of sediment within a narrower, but thicker on-lapping sediment prism. thick back-barrier deposits accumulated below the ravinement surface or tsme, while a sheet of transgressive shelf sands was shed seawards. after infill of the remaining accommodation space with hst/fsst deposits, re-establishment of coal-forming mires and swamps indicates a new tectonically quiescent phase. thus periods characterised by tectonic quiescence and slow uniform subsidence alternated with episodes of faulting at the main boundary fault, when the hangingwall slope was re-established and the newly created accommodation space was filled. however, compaction of thick peat may also have favoured the preferential preservation of transgressive deposits in the western part of the søgne basin where the thickest coals are found. similar relationships have been described by ravnås & steel (1998) from the middle jurassic tarbert formation in the northern north sea. these workers described how the destruction of shoreline barriers by steep-trajectory transgression resulted in sediment being partitioned landwards and seawards. both the overall, gradual change from alluvial plain or fluvially dominated coastal plain deposits in the lower part to dominantly tidal and shallow marine deposits in the upper part of the middle jurassic succession and the backstepping stacking pattern of the uppermost three sequences indicate that not only punctuated rift-related subsidence but also a large-scale eustatic sea-level rise or regional subsidence participated in the creation of accommodation space. the important sequence boundary at the base of the cal-1a sequence, which formed in late bathonian or earliest callovian times, cuts deeply into deposits both on the upper hangingwall slope and in basinal locations close to the main boundary fault. this sequence boundary can be traced into the southern part of the danish central graben (michelsen et al. 2003, this volume, fig. 36). this supports the suggestion that a regional fall in relative sea level rather than local rift-related tectonics is responsible for the formation of that sequence boundary. as a result of the tectonic influence on sedimentation in the latter part of the middle jurassic, both estuarine and shoreface depositional systems, which may both contain important reservoir rocks, show a systematic distribution pattern that is related to the halfgraben geometry of the basin and therefore potentially predictable. shore and shoreface sandstones of the hst and fsst in the uppermost sequences occur as strikeparallel laterally extensive sheet sandstones. they can be correlated with negligible changes in thickness, grain size and facies for at least 15 km in the danish søgne basin. in contrast, thick estuarine channel deposits of the tst in the uppermost sequences mainly occur in dip-parallel incised valleys. 344 acknowledgements this study formed part of a ph.d. undertaken at copenhagen university. i am grateful to my supervisor finn surlyk for his thorough constructive critisism, which contributed significantly to the improvement of this paper, to reviewers jan alexander and guy plint for their helpful comments and to jon ineson for thorough editing. the work was supported by efp-92 grant no. 1313/92-0002 from the danish energy agency and by mærsk oil and gas a/s and norsk hydro udforskning a/s. i had fruitful discussions with colleagues karen dybkjær, jon r. ineson, peter johannessen and lars h. nielsen. karen dybkjær and niels poulsen kindly supplied me with palynological datings. references alexander, j. & gawthorpe, r.l. 1993: the complex nature of a jurassic multi-storey 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(eds): tectonic evolution of the north sea rifts, 1–36. oxford: clarendon press. 347 manuscript received 30 may 1997; revision accepted 19 december 2001. bat-1b sb bat-1b mfs cal-1a sb 5 m m m m 4504 4511 4530 3618 3639 3680 3509 3540 3574 3/7-4 6 km 4.5 km gr west lulu-1 gr lulita-1 gr crevasse delta crevasse splay fluvial channels fluvial channels lake and distal floodplain fluvial channels ■ ■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ amalie-1 3/7-4 lulita-1 lulu-1 west lulu 1 3 2 4 5l coal claystone siltstone sandstone lithology depositional environment floodplain fluvial channels key surfaces sequence boundary (sb) maximum flooding surface (mfs) siclay sand gr siclay sand gr siclay sand gr søgne basin 10 km fig. 15. log panel (gr and core logs) depicting floodplain deposits with channel and crevasse sandstones and lacustrine mudstones of sequence bat-1b in the 3/7-4, west lulu-1 and lulita-1 wells. thick lacustrine mudstones are located around the bat-1b mfs in west lulu-1. the succession is incised by the cal-1a sb, marking the base of an incised valley. the positions of cores illustrated with photographs here are indicated on the sedimentological logs (e.g. 5l indicates core photograph in fig. 5l). depths of important surfaces, facies changes or core breaks are indicated (in metres below reference level). for full legend, see fig. 13; inset map shows the location of the transect. cal-1a lulu fm bryne fm cal-1a sb west lulu-2 gr 2.5 km 1.8 km 1.4 km west lulu-4 gr west lulu-3 gr west lulu-1 gr 3668 m m m m 3666 3685 3705 3711 6l 6k 6j 6i 6h 6g 6f 6e 6d 6c 6b 6a 3592 3606 3618 3834 3842 3848 3681 coastal mire coastal mire passive channel fill stacked estuary channels fluvial stacked estuary channels stacked estuary channels transgressive reworked barrier bay/lagoon stacked estuary channels ■ ■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ amalie-1 3/7-4 lulita-1 lulu-1 west lulu 1 3 2 4 coal claystone siltstone sandstone conglomerate floodplain mires, swamps lagoon and tidal flats estuary and tidal channels, bay-head deltas shoreface, mouth bars and washovers sequence boundary (sb) maximum flooding surface (mfs) lithology depositional environments key surfaces siclay sand gr siclay sand gr siclay sand gr siclay sand gr 5 m søgne basin 10 km fig. 16. log panel (gr and core logs) showing incised valley fill deposits from the uppermost bryne formation (lst/lower tst of sequence cal-1a). possible lst deposits are limited to the lowermost two metres of the succession in west lulu-3. most channel sandstones show abundant sedimentary structures indicating a tidally influenced environment and represent estuary channel deposits. line of section is broadly sw–ne, perpendicular to the inferred valley axis (see inset map). the positions of cores illustrated with photographs here are indicated on the sedimentological logs (e.g. 6a indicates core photograph in fig. 6a). depths of important surfaces, facies changes or core breaks are indicated (in metres below reference level). for full legend, see fig. 13. amalie-1 gr estuary channel estuary channel tidal flat stacked estuary channels stacked estuary channels lagoon/tidal flats bay/lagoon estuary channels estuary channels swamps estuary channels tidal channels and swamps outer or marginal estuary with minor channels lagoon and tidal flats bryne fm cal-1a sb lulu fm ■ ■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ amalie-1 3/7-4 lulita-1 lulu-1 west lulu 1 3 2 4 3476 m 3488 3509 4504 4480 5071 m 5088 5110 5119 4453 m 8a 8f 8c 8d 8e 8b 3/7-4 2 km 14 km gr lulita-1 gr coal claystone siltstone sandstone conglomerate floodplain mires, swamps lagoon and tidal flats estuary and tidal channels, bay-head deltas marine shelf sequence boundary (sb) maximum flooding surface (mfs) lithology depositional environments key surfaces siclay sand gr siclay sand gr siclay sand gr 5 m søgne basin 10 km fig. 17. log panel (gr and core logs) showing incised valley fill and valley mouth deposits (tst of sequence cal-1a). wells 3/7-4 and lulita-1 represent distal valley fill or valley mouth deposits of the incised valley also depicted in fig. 17. amalie-1 represents valley fill deposits from an incised valley in the southern part of the søgne basin. the positions of cores illustrated with photographs here are indicated on the sedimentological logs (e.g. 8b indicates core photograph in fig. 8b). depths of important surfaces, facies changes or core breaks are indicated (in metres below reference level). for full legend, see fig. 13; inset map shows the roughly nnw–sse trend of the transect, broadly axial in the søgne basin. ? ? ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ amalie-1 3/7-4 lulita-1 lulu-1 west lulu 1 3 2 4 west lulu-1 gr 3/7-4 gr west lulu-2 2.1 km 1.4 km 6 km gr west lulu-3 gr 10a 10b 3780 m 3617 m 3625 3636 3646 3666 3592 3575 3572 3567 3560 m 3440 m 3449 3459 3460 3464 3476 3799 3800 3808 10c 10d 10e 10f 10g 10i 7a 7b 7c 9a 9b 9c 9d 9e 9f 9g 9h 9i 10h 12a 12b 12c lagoon estuary channels and bars lagoon bay-head delta mouth bar offshore lola fm lulu fm shoreface shoreface or mouth bar bay-head delta mouth bar washover and ravinement complex mires bay/lagoon lagoon bay-head delta with distributary channels estuary channels and bars cal-1c cal-1c sb cal-1b cal-1b sb cal-1a lola fm upper paralic wedge upper marine wedge middle paralic wedge lower paralic wedge lower marine wedge cal-1c sb cal-1b sb cal-1b mfs rsme cal-1a mfs lulu fm lulu fm bryne fm ■ ■ ■ ■ ■ ■ coal claystone siltstone sandstone conglomerate floodplain mires, swamps lagoon and tidal flats estuary and tidal channels, bay-head deltas shoreface, mouth bars and washovers marine shelf sequence boundary (sb) maximum flooding surface (mfs) regressive surface of marine erosion (rsme) normal fault lithology depositional environments key surfaces siclay sand gr siclay sand gr siclay sand gr siclay sand gr 5 m søgne basin 10 km fig. 18. log panel (gr and core logs) showing paralic and marine wedges of the lulu formation (upper tst and hst of cal-1a, cal-1b and lst and lower tst of cal-1c). the transect (see inset map) is largely within the westernmost part of the basin, dominated by the paralic sediment wedges. the marine wedges are present only in the west lulu-1 and 3/7-4 wells, wedging out between paralic wedges to the west. the positions of cores illustrated with photographs here are indicated on the sedimentological logs (e.g. 10a indicates core photograph in fig. 10a). depths of important surfaces, facies changes or core breaks are indicated (in metres below reference level). for full legend, see fig. 13. shelf shelf transgressive shelf transgressive shelf prograding shoreface prograding shoreface prograding shoreface or mouth bar shoreface or mouth bar back barrier swamps swamps shoreface offshore offshore shoreface ■ ■ ■ ■ ■ ■ amalie-1 gr lulu-1 11 km gr lulita-1 3.5 km gr 3/7-4 2 km gr 3440 m 3449 3459 3460 3464 4454 4448 4439 4434 4426 m 12f 12g 12h 12i 12j 11b 11c 11d 11e 11f 11g 11h 3574 m 3584 3587 3595 11a 12e 12d 3600 5071 5061 5048 m 3476 lola fm lulu fm lulu fm bryne fm cal-1c cal-1b cal-1a cal-1c sb cal-1b sb cal-1b mfs rsme cal-1a mfs lower marine wedge upper marine wedge ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ amalie-1 3/7-4 lulita-1 lulu-1 west lulu 1 3 2 4 coal claystone siltstone sandstone floodplain mires, swamps lagoon and tidal flats estuary and tidal channels, bay-head deltas shoreface, mouth bars and washovers marine shelf sequence boundary (sb) maximum flooding surface (mfs) lithology depositional environments key surfaces siclay sand gr siclay sand gr siclay sand gr siclay sand gr 5 m søgne basin 10 km fig. 19. log panel (gr and core logs) showing marine and paralic wedges of the lulu formation (upper tst and hst of cal1a, cal-1b and lst and lower tst of cal-1c). line of section is nnw–sse (see inset map) roughly parallel to the palaeocoastline in the central part of the basin and dominated by deposits of the marine sedimentary wedges. note that the prograding shoreface packet in the upper levels of the cal-1c sequence has an abrupt erosional base; this surface is interpreted as a regressive surface of marine erosion (rsme), defining the base of the falling stage systems tract. the positions of cores illustrated with photographs here are indicated on the sedimentological logs (e.g. 12f indicates core photograph in fig. 12f). depths of important surfaces, facies changes or core breaks are indicated (in metres below reference level). for full legend, see fig. 13. geological survey of denmark and greenland bulletin 26, 2012, 9-12 9 nano-quartz in north sea danian chalk holger lindgreen and finn jakobsen the main oil reservoir in the central graben in the north sea is chalk of the maastrichtian tor formation, which has high porosity and relatively high permeability. the chalk of the danian ekofisk formation is an additional reservoir, but with highly variable porosity and permeability. whereas the tor formation is almost pure calcite primarily consisting of coccolith debris, the ekofisk formation also comprises significant proportions of phyllosilicates (clay minerals) and quartz in addition to coccolith debris. for decades the quartz was assumed to be a normal crystalline α-quartz such as is present in quartz sand, and the clay fraction was assumed to consist predominantly of phyllosilicates. however, maliva & dickson (1992) reported the presence of presumably authigenic submicron-size quartz crystals arranged in clusters, and suggested that these clusters were transformed opal-ct lepispheres. investigations by nano-structural methods (xray diffraction and atomic force microscopy (afm)) revealed that the prevailing quartz component in the north sea chalk comprises α-quartz appearing as nano-size quartz spheres (jakobsen et al. 2000; lindgreen et al. 2010). nano-quartz spheres were first observed in indurated chalk in the ekofisk formation in the ekofisk field and later in the south arne field. subsequent analyses of the ekofisk formation in different chalk fields showed that the content of nano-quartz varies throughout the chalk succession and to some degree reflects the cyclic development of the chalk. the proportion of dispersed nano-quartz in the chalk is highly variable, from 10% to more than 80% in the lower danian (lindgreen et al. 2010). this paper describes the nano-quartz, its formation and structure and presents a model for the formation of flint from nano-quartz in the north sea ekofisk chalk. material and methods we have investigated core samples from the ekofisk formation in the south arne field wells sa-1 and rigs-1, in the halfdan field wells sif-1 and nana-1 and in the gorm field well n-22 (fig. 1). most samples contained large amounts of calcite, so calcite-free residues were prepared by dissolving the calcite in an acetate-acetic acid buffer at ph 4.5–5. in this buffer, non-calcite minerals and especially fine-grained nano-quartz and clay minerals are not corroded or dissolved. scanning electron microscopy (sem) is routinely used for investigations of chalk minerals and in special cases micronsized particles can be identified (hjuler & fabricius 2009). however, rock samples dominated by nano-sized quartz are at the limit of resolution in the sem and generally give poor sem images due to poor current transmission in the fine-grained matrix. we used x-ray diffraction and afm to characterise the ultra-fine particles in the chalk, such as nano-quartz and clay minerals. x-ray diffraction scanning using 10 s/0.1 °2θ was applied routinely to determine the mineralogical composition of both core piece samples and of non-calcite residues. high statistic scanning using 100 s/0.02 °2θ was used to characterise the nano-quartz. afm (binnig et al. 1986) generates topographic images by scanning a sharp tip across a surface and can produce images at atomic resolution of both conductors and nonconductors. for afm we used a rasterscope 3000 instrument under room conditions with a force of 0.175 nn and a scanning speed of 500 nm/s. in the present investigation of the topography of raw surfaces, afm was run in non-contact r ingkøbing–fyn h igh gas in chalk oil in chalk field at other level 4°e 5°e 56°n c entral g raben 25 km national border rigs-1 sa-1 sif-1 nana-1 n-22 uk nl n dk 200 km g fig. 1. map of the danish central graben showing the locations of the investigated wells. © 2012 geus. geological survey of denmark and greenland bulletin 26, 9–12. open access: www.geus.dk/publications/bull 1010 mode. intact rock samples of small pieces of chalk or flint were glued onto gold-coated sample holders. in chalk samples, non-calcite minerals were identified and imaged from the insoluble residue. for such samples, the residue was dispersed ultrasonically in distilled water and the samples prepared by leaving a drop of the suspension to dry under room conditions on a block of highly oriented pyrolytic graphite. structure of the nano-quartz particles afm of non-calcite residues deposited on graphite showed that the nano-quartz consists predominantly of rather uniform, c. 600 å large spherical particles (fig. 2a). afm images of intact flint surfaces showed that the flint consists of similar spherical particles with a diameter of c. 500 å or more (fig. 2b) and some irregularly shaped particles. x-ray diffraction showed that the non-calcite residues and the flint and quartz layers are composed of α-quartz having practically identical patterns and resembling the pattern of standard quartz (fig. 3). it is remarkable that the quartz in all the examined samples of dispersed quartz and flint have almost identical unit cell a and c parameters and sizes of coherent scattering domains (lindgreen et al. 2011). at high angles peak broadening was pronounced for the nano-quartz particles and careful recording revealed a broad and distorted pattern of the (212), (203) and (301) reflections compared to the reflection from standard quartz (fig. 4). these distortions are due to larger a and c parameters compared to those of normal quartz (lindgreen et al. 2011). the nano-quartz spheres had colloidal properties and flocculated in suspensions with sufficient ionic strength, such as sea water (fig. 5). formation of quartz particles the nano-quartz spheres are anticipated to be of a type that might crystallise in a marine environment which is slightly enriched in silicon (williams & crerar 1985). the source of silicon was probably opal-a from radiolarians, which were the main silica-bearing organism in the chalk sea (maliva & dickson 1992). it is important that the non-crystalline si in radiolarians will dissolve at the low concentration of si, which is sufficient to precipitate fine quartz, and that the quartz will be the first silica phase to crystallise. a b 200 nm200 nm standard merck quartz sa 3344.15 m quartz in matrix nana 2135.7 m quartz in flint 80 81 °2θ 82 2 1 2 α 1 2 0 3 α 1 2 0 3 α 2 3 0 1 α 1 3 0 1 α 2 2 1 2 α 2 a b c fig. 3. a: x-ray diffraction pattern of standard merck quartz. b: of calcitefree residue from chalk, south arne field, well sa-1, 3344.15 m, c: and of flint layer in halfdan field, well nana-1, 2135.7 m. co-kα radiation, 5% si added as internal standard. fig. 4. x-ray diffraction patterns. a: region of (212), (203) and (301) reflections of standard merck quartz. b: the same region for calcite-free residue from chalk, south arne field, well sa-1, 3344.15 m. c: the same region for flint layer in halfdan field, well nana-1, 2135.7 m. co-kα radiation. fig. 2. atomic force microscopy images of spherical grains of quartz. noncontact mode, room conditions, force 0.175 nn, scanning speed 500 nm/s. a: calcite-free residue deposited on graphite from well sa-1, 3344.15 m. b: intact flint from nana-1, 2135.7 m. 20 30 40 50 60 70 80 90 standard merck quartz sa 3344.15 m quartz in matrix nana 2135.7 m quartz in flint si si si °2θ a b c 11 data from williams et al. (1985) indicate that the c. 500 å diameter quartz spheres observed in the chalk and in the flint of the north sea danian chalk can form at sio2 concentrations of c. 12 ppm. the north sea chalk is a deep water deposit and present-day deep sea water has a concentration of 1–10 ppm sio2 (millot 1970; calvert 1974). we think that only a minor increase in si concentration would result in crystallisation of nano-quartz spheres. the colloidal quartz spheres could then have flocculated and been deposited on the sea floor mixed with coccolith ooze. flocculation is important for sedimentation of silica and the rate of sedimentation for the formation of layers rich in quartz. chemical environment in the water column as described above, we assume that silica was not deposited as biogenic opal-a. therefore the variation in proportion of nano-quartz cannot be caused by changes in the supply of silicon to the sea as such changes would be reflected in changes in size and mineralogy of the silica. an alternative is variation in the sedimentation of coccoliths. such variation may be due to a decrease in ph which may cause coccoliths to be partly or totally dissolved in the water column. such a decrease in ph requires significant amounts of an acidifier. this acidifying agent was most probably atmospheric co2, which by mixing with sea water has been found to decrease the calcification of marine plankton (riebesell et al. 2000; feely et al. 2004). co2 released in large quantities during volcanic eruptions (holmes 1965; zimmer & erzinger 2003; frondini et al. 2004; schuiling 2004; self et al. 2006) could be a cause of the dissolution of the coccoliths in parts of the danian chalk deposits in the north sea. sensitivity analysis has indicated that only massive and short-lived volcanism could cause the caco3 undersaturation of seawater (berner & beerling 2007). age determinations of lavas from the british tertiary igneous province have yielded ages of 63–65 ma (saunders et al. 1997), corresponding to a danian age. we therefore propose that the pronounced quartz enrichment in the danian chalk of the north sea was associated with frequent volcanic eruptions in this period at and after the cretaceous–tertiary boundary. our model implicates that the degree of dissolution of the coccoliths in the sedimentary environment determines the proportion between calcite and nano-quartz in the chalk. theories for flint formation the new theory for the formation of flint and dispersed nano-quartz in the north sea by crystallisation of nano-quartz in the marine environment is totally different from the generally accepted theory for flint formation in chalk based on studies of chalk from onshore outcrops (bromley & ekdale 1986; clayton 1986; zijlstra 1987; madsen & stemmerik 2010). according to the current theory for flint formation, opaline tests and sponge spicules in the sediment are dissolved during burial and the si is recrystallised as opal-ct and quartz in hollows and by replacement of calcite. however, the generally accepted theory does not agree with our results obtained for the silica in the investigated north sea chalk. we support our point of view by observing that the type of α-quartz dispersed in chalk is identical with the type constituting the flint nodules and flint horizons and with the type constituting the α-quartz horizons in the chalk. we find it highly unlikely that the same size and shape of particles will crystallise in the chalk and in the hollows during dissolution or reprecipitation, whereas the particles are of the type which can possibly crystallise in the marine environment that is slightly enriched in silicon (williams & crerar 1985). a sedimentary origin of the silica-rich chalk layers is supported by the presence of a flint bed in well n-22. the flint layer includes a calcite-filled burrow within fig. 5. flocculation of nano-quartz particles. residue from well sa-1, 3353.0 m. left: quartz dispersed in distilled water. right: quartz dispersed in 0.2 m cacl2. dispersed in distilled water dispersed in 0.2m cacl2 1212 a matrix comprising nano-quartz spheres. the occurrence of a burrow in rather pure α-quartz sediment shows that the α-quartz was soft when biological activity took place. a sedimentary origin of the flint fits well with our results for the north sea tertiary chalk, which is a deposit in relatively deep water. however, it cannot be generally applied to other areas and deposits in different settings without further investigations. conclusions we have proposed a new model for the formation of flint in north sea chalk: (1) the nano-quartz in the flint, like the nano-quartz in the chalk matrix, has crystallised in the marine chalk-sea environment. the colloidal quartz particles have then flocculated and have been deposited on the sea floor mixed with coccolith ooze. (2) regional variations in the concentration of nano-quartz particles in the sediment reflect different degrees of acidification of the chalk sea. (3) this resulted in areas with a high concentration of nanoquartz spheres that could form flint layers. in areas with lower concentration of nano-quartz spheres, indurated chalk with abundant nano-quartz particles are now preserved. (4) the acidification may have been caused by enhanced atmospheric co2 linked to massive volcanic eruptions. references berner r.a. & beerling d.j. 2007: volcanic degassing necessary to produce a caco3 undersaturated ocean at the triassic–jurassic boundary. palaeogeography, palaeoclimatology, palaeoecology 244, 368–373. binnig, g., quate, c.f. & gerber, ch. 1986: atomic force microscope. physics revue letters 56, 930–933. bromley r.g. & ekdale, a.a. 1986: flint and fabric in the european chalk. in: sieveking, g.d.g. & hart, m.b. (eds): the scientific study of flint and chert, 71–82. cambridge: cambridge university press. calvert s.e. 1974: deposition and diagenesis of silica in marine sediments. international association of sedimentologists, special publication 1, 273–300. clayton c.j. 1986: the chemical environment of flint formation in upper cretaceous chalk. in: sieveking, g.d.g. & hart, m.b. (eds): the scientific study of flint and chert, 43–54. cambridge: cambridge university press. feely, r.a., sabine c.l., lee, k., berelson, w., kleypas, j., fabry, v.j. & millero, f. j. 2004: impact of anthropogenic co2 on the caco3 system in the oceans. science 305, 362–366. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: hl@geus.dk frondini, f., chiodini, g., caliro, s., cardellini, c., granieri, d. & ventura, g. 2004: diffuse co2 degassing at vesuvio, italy. bulletin of volcanology 66, 642–651. hjuler, m.l. & fabricius, i.l. 2009: engineering properties of chalk related to diagenetic variations of upper cretaceous onshore and offshore chalk in the north sea area. journal of petroleum science and engineering 68, 151-170. holmes, a. 1965: principles of physical geology, 1288 pp. london: nelson. jakobsen, f., lindgreen, h. & springer, n. 2000: precipitation and flocculation of spherical nano silica in north sea chalk. clay minerals 35, 175–184. lindgreen, h., jakobsen, f. & springer, n. 2010: nano-size quartz accumulation in reservoir chalk, ekofisk formation, south arne field, north sea. clay minerals 45, 171–182. lindgreen, h., drits, v.a., salyn, a.l., jakobsen, f. & springer, n. 2011: formation of flint horizons in north sea chalk through marine sedimentation of nano-quartz. clay minerals 46, 525-537. madsen, h.b. & stemmerik, l. 2010: diagenesis of flint and porcellanite in the maastrichtian chalk at stevns klint, denmark. journal of sedimentary research 80, 578–588. maliva, r.g. & dickson, j.a.d. 1992: microfacies and diagenetic controls of porosity in cretaceous/tertiary chalks, eldfisk field, norwegian north sea. aapg bulletin 76, 1825–1838. millot, g. 1970: geology of clays, 429 pp. new york: springer. riebesell, u., zondervan, i., rost, b., tortell, p.d., zeebe, r.e. & morel, f.m.m. 2000: reduced calcification of marine plankton in response to increase atmospheric co2. nature 407, 364–367. saunders, a.d., fitton, j.g., kerr, a.c., norry, m.j. & kent, r.w. 1997: the north atlantic igneous province. in: mahoney, j.j. & coffin, m.f. (eds): large igneous provinces. geophysical monograph series 100, 45–97. schuiling, r.d. 2004: thermal effects of massive co2 emissions associated with subduction volcanism. comptes rendus geoscience 336, 1053–1059. self, s., widdowson, m., thordarson, t. & jay, a.e. 2006: volatile fluxes during flood basalt eruptions and potential effects on the global environment: a deccan perspective. earth and planetary science letters 248, 518–532. williams, l.a. & crerar, d.a. 1985: silica diagenesis, ii. general mechanisms. journal of sedimentary petrology 55, 312–321. williams, l.a., parks, g.a. & crerar, d.a. 1985: silica diagenesis, i. solubility controls. journal of sedimentary petrology 55, 301–311. zijlstra, h.j.p. 1987: early diagenetic silica precipitation, in relation to redox boundaries and bacterial metabolism, in late cretaceous chalk of the maastrichtian type locality. geologie en mijnbouw 66, 343–355. zimmer, m. & erzinger, j. 2003: continuous h2o, co2, 222rn and temperature measurements on merapi volcano, indonesia. journal of volcanology and geothermal research 125, 25–38. research article | short bennike et al. 2020: geus bulletin 44. 4630. https://doi.org/10.34194/geusb.v44.4630 1 of 6 late quaternary history of lammefjorden, north-west sjælland, denmark ole bennike*1, peter roll jakobsen1, jakob walløe hansen2 1geological survey of denmark and greenland (geus), aarhus, denmark. 2unesco global geopark odsherred, nykøbing, denmark. abstract lammefjorden is a reclaimed fjord in north-west sjælland, denmark. sediment cores from the area were collected to study its development after the last deglaciation, in particular the sea-level history. late glacial and early holocene lake and bog deposits occur below marine deposits. sparse late glacial fossil assemblages indicate tree-less environments with dwarf-shrub heaths. early holocene deposits contain remains of betula sec. albae sp. and pinus sylvestris, which indicate open forests. the wetland flora comprised the calciphilous reed plant cladium mariscus and the water plant najas marina. marine gyttja from basins is characterised by sparse benthic faunas, probably due to high sedimentation rates. in some areas, shell-rich deposits were found, with large shells of ostrea edulis, indicative of high summer temperatures, high salinity and strong tidal currents. a marine shell dated to 6.7 cal. ka provides a minimum age for the marine transgression of lammefjorden. introduction in 1873, a large project was initiated to reclaim lammefjorden in north-west sjælland (fig. 1), and today lammefjorden is one of the largest reclaimed areas in north-west europe. the geology of the lammefjorden region was mapped by the geological survey of denmark in the late 1890s (rørdam & milthers 1900). shell samples from holocene-raised marine deposits were analysed and several of them contained shells of ostrea and tapes. these bivalves no longer live in the region and their former presence was taken to indicate stronger tidal currents than at present (rørdam & milthers 1900, p. 101). it was also noted that in situ quercus stumps and peat deposits were found below marine deposits. both marine and non-marine deposits were referred to as ‘alluvium’ (holocene), but a more precise age could not be determined. madsen et al. (1900) described a large stone age shell midden near fårevejle in the inner part of lammefjorden (fig. 1). the mollusc fauna comprised shells of ostrea and tapes. the ostrea shells were up to 130 mm long – a large size pointing to favourable conditions. artefacts mainly belonged to the late mesolithic ertebølle culture. the shell midden was located at an elevation of 3.5 to 5 m a.s.l., just above the marine limit at c. 3.5 m a.s.l. (mertz 1924). in the 1920s and 1930s, archaeologist erik westerby tried to locate submarine, pre-ertebølle settlements in denmark. westerby (1933) searched for such sites in reclaimed areas now situated below sea level and found *correspondence: obe@geus.dk received: 19 feb 2020 accepted: 23 apr 2020 published: 25 june 2020 keywords: lammefjorden, sjælland, quaternary, late glacial, holocene, macrofossils abbreviations: a.s.l.: above sea level b.s.l.: below sea level rsl: relative sea level unesco: united nations educational, scientific and cultural organization geus bulletin is an open access, peerreviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: catherine jex (geus, denmark) reviewed by: jonathan lewis (loughborough university, uk) and one anonymous reviewer funding: see page 5 competing interests: none declared additional files: see page 5 https://doi.org/10.34194/geusb.v44.4630 mailto:obe@geus.dk https://creativecommons.org/licenses/by/4.0/deed.ast bennike et al. 2020: geus bulletin 44. 4630. https://doi.org/10.34194/geusb.v44.4630 2 of 6 www.geusbul let in.org settlements at two sites in lammefjorden. they belong to what is now referred to as the kongemose culture. one of the sites near gislinge (fig. 1), located at 2.5 m b.s.l., was excavated by mathiassen (troels-smith 1942) and by simonsen (1946). christensen (1994) and christensen & andreasen (1999) discussed rsl changes in the region. dating of charcoal from the gislinge settlement places it in the early kongemose culture (k-5383, supplementary file s1) whereas charcoal from storø (fig. 1) indicates a mid-kongemose culture age (k-5324, supplementary file s1). these data acquired by archaeologists provided the first information on holocene sealevel changes. in 2014, geopark odsherred was established as a unesco global geopark (jakobsen et al. 2015). the inner part of lammefjorden is included in the geopark. to further develop the record of holocene sea-level change of the former fjord using modern techniques, four sediment cores from the area were collected (cores 197.670–73) in 2016, and in 2017 other seven cores were collected, including 197.77 and 197.79 reported here (fig. 1). in 2017, georadar data were also collected. this work was reported in unpublished (non-peer reviewed) reports in danish by jakobsen et al. (2016) and jakobsen & bennike (2017). here, the results for an international audience are presented, along with new descriptions and sampling of open sections studied in 2019 (sites a/b in fig. 1). in total, 86 samples were analysed for macrofossils, and 12 samples of macrofossils were submitted for radiocarbon dating. the sampling locations are shown in fig. 1 and methods are provided in an online supplementary file. results long cores simplified sedimentological core logs for six cores are shown in fig. 2, and macrofossil data are shown in supplementary file s2. core sites were chosen to provide long records. core 197.670 was 13.8 m long and consisted of clayey till overlaid by two units of glaciofluvial sand. the upper sand unit contained rare remains of, for example, dryas octopetala, betula nana and distichium sp., which indicate an environment with dwarf-shrub heaths of late glacial age. a similar fossil assemblage from core 197.671 was dated to 13.9 cal. ka (see below). the late glacial sediments are overlaid by marine and brackish-water gyttja with shells of, for example, mytilus edulis, cerastoderma sp., hydrobia sp., bittium reticulatum, balanus crenatus and cyprideis torosa (ostracode). the faunas indicate a marine environment with moderate salinity. mytilus and balanus need firm substrates. a sample from the base of the marine mud was dated to 2.5 cal. ka. if sedimentation continued till the area was reclaimed, it means that 10 m of gyttja was deposited in 2.5 ka, corresponding to a high mean sedimentation rate of 4 mm/year. with only one date, this sedimentation rate is obviously highly uncertain. core 197.671 was 25 m long and consisted of till overlain by clay, silt and sand with abundant shells of the small crustacean alona sp. and rare remains of chironomids and cristatella mucedo. these species indicate a lacustrine environment. land plants were mainly represented by betula sect. albae. one sample from 26.75 m b.s.l. contained remains of betula nana, dryas octopetala, distichium sp., hippuris vulgaris and aulacomium palustre. fig. 1 map of north-west sjælland showing the reclaimed area of lammefjorden; 670–679 denote core sites for cores 197.670 to 197.679. a/b: open sections sampled in 2019. gr: georadar profile 1, see fig. 3 for data. få: fårevejle. to: toftevang. st: storø. gi: gislinge lammefjorden. https://doi.org/10.34194/geusb.v44.4630 http://www.geusbulletin.org bennike et al. 2020: geus bulletin 44. 4630. https://doi.org/10.34194/geusb.v44.4630 3 of 6 www.geusbul let in.org a sample from this level was dated to 13.9 cal. ka. alona indicates lake environment and land plants indicate late glacial age (11.7–14.7 cal. ka), as confirmed by 14c dating. the lake sediments are overlaid by marine gyttja with shells of mytilus, cerastoderma, hydrobia sp., bittium reticulatum and cyprideis torosa as well as rare remains of littorina littorea, rissoa parva, tritia reticulata and balanus crenatus. core 197.672 was 9 m long, with clayey till at the base and 7.4 m marine gyttja with mytilus, cerastoderma, scrobicularia plana, corbula gibba, kurtiella bidentata, hydrobia sp., balanus crenatus and cyprideis torosa. a shell sample near the base yielded an age of 2.8 cal. ka, indicating a mean sedimentation rate of 2.4 mm/year. core 197.673 was 10 m long. the lower part of the core contained till and glaciofluvial sediments with in situ roots of land plants. the upper part consisted of gyttja with macrofossils that comprise mytilus, ostrea, cerastoderma, hydrobia sp., bittium reticulatum and cyprideis torosa. core 197.677 was 20.5 m long. the core contained sandy silt with alona sp. at the very bottom, interpreted as late glacial lake sediment. it is followed by 20.4 m marine or brackish-water gyttja. the lower part contained only scattered, rare small fragments of mytilus, a few jaws of nereis sp. and some skeletal remains of hydroids. shells become more frequent at c. 15.5 m b.s.l. and the upper half of the core comprised mytilus, cerastoderma, littorina littorea, hydrobia sp., bittium reticulatum and rare retusa truncatula. a sample from 15.6 m b.s.l. is dated to 2.0 cal. ka, indicating a very high mean sedimentation rate of nearly 6 cm/year. with only one date, this sedimentation rate is obviously highly uncertain as mentioned above. georadar data and a short 2 m core georadar profiling was carried out at two sites with oyster banks (fig. 3). at the first site (profile 1), the elevation increases from 4 m b.s.l. in the west to 1.3 m b.s.l. in the east. in the east, with no penetration, clayey till was found by coring. the eastern part of the profile was followed by a 20 m long section with strong reflectors. coring revealed alternating layers of shells and sand, followed by gravelly sand and clayey till. in the western part of the profile, the penetration depth of the radar signal was only c. 1 m and the reflectors were weak. the internal structures show progradation to the west. coring showed fine-grained sand, silt and gyttja with marine shells in this area. at the second site, profile 2, numerous marine shells occurred on the terrain surface. in the westernmost part of the profile, with almost no penetration, coring revealed clayey till. the rest of the profile showed relatively strong reflectors, which mark layering with two shell banks. core 197.679 was collected in the western structure, it was 2 m long with clayey till at the base, followed by organic sediments with shells of terrestrial snails (carychium tridentata, vallonia costata, vertigo pusilla and vitrea contracta; fig. 2 and supplementary file s2). this was followed by sand and gravel with shells of fig. 2 simplified sedimentological core logs from lammefjorden. further details on sediments are available at http://www.geus.dk/ produkter-ydelser-og-faciliteter/ data-og-kort/national-boringsdatabase-jupiter/. the ages show median calibrated ages in years before the present (bp). https://doi.org/10.34194/geusb.v44.4630 http://www.geusbulletin.org http://www.geus.dk/produkter-ydelser-og-faciliteter/data-og-kort/national-boringsdatabase-jupiter/ http://www.geus.dk/produkter-ydelser-og-faciliteter/data-og-kort/national-boringsdatabase-jupiter/ http://www.geus.dk/produkter-ydelser-og-faciliteter/data-og-kort/national-boringsdatabase-jupiter/ http://www.geus.dk/produkter-ydelser-og-faciliteter/data-og-kort/national-boringsdatabase-jupiter/ bennike et al. 2020: geus bulletin 44. 4630. https://doi.org/10.34194/geusb.v44.4630 4 of 6 www.geusbul let in.org marine species. the fauna comprised abundant shells of mytilus and balanus crenatus, common shells of ostrea edulis and bittium reticulatum and rare shells of, for example, cerastoderma sp., tritia reticulata and buccinum undatum. two samples were dated, one at the base of the shell-rich bed and another at the top. the ages of the samples were c. 4.5 and 3.1 cal. ka. open sections two open sections were studied in 2019 (sections a and b; fig. 1), located 400 m apart. the results are summarised in supplementary files s3 and s4. the lowermost sediments at site a (7 m b.s.l.) consisted of lake marl with shells of fresh-water molluscs such as valvata sp., bithynia tentaculata, radix peregra, acroloxus palustris and pisidium sp., statoblasts of cristatella mucedo. remains of the water plants chara sp., nymphaea alba and potamogeton natans, the reed plants phragmites australis and cladium mariscus and the land plants betula sect. albae and pinus sylvestris were also found. two samples were dated to c. 10.3 and 10.6 cal. ka (supplementary file s1). the marl was overlain by coarse-grained detritus gyttja at 6 m b.s.l. with remains of the water plants nymphaea alba, potamogeton natans and najas marina, the reed plants schoenoplectus lacustris and cladium mariscus, and the land plants betula sect. albae and pinus sylvestris. a sample from this layer was dated to 8.8 cal. ka. najas marina and cladium mariscus are rare in denmark today, but they were common in the early holocene (bennike et al. 2004). the lake sediments were overlaid by a layer rich in shells of marine molluscs at 5.5 m b.s.l.. the fauna included ostrea, mytilus, bittium reticulata and tritia reticulata. the largest ostrea shell was 105 mm long and dated to c. 6.2 cal. ka. finally, a non-in situ root of pinus from 6.5 m b.s.l. from the section was dated to 9.9 cal. ka. at site b, a sample of forest peat from 7 m b.s.l. mainly consisted of twigs and wood fragments with many bark fragments of pinus sylvestris and a few sclerotia of the fungus cenococcum geophilum (supplementary file s3). a sample of pinus sylvestris bark fragments yielded an age of c. 9.8 cal. ka (supplementary file s1). discussion based on the cores described here, it is clear that there was an at least 20 m deep basin in the inner western part of lammefjorden. we suggest that it formed as glacier bulldozed material into the arc-shaped hilly area west of lammefjorden, as first proposed by rørdam & milthers (1900). sjælland was de-glaciated about 18  000 to 16  000 years ago (houmark-nielsen et al. 2012), but it apparently took several millennia before plants immigrated to the region. we obtained an age of c. 13.9 cal. ka – the oldest plant remains so far dated from the region. however, older samples, up to c. 15.2 cal. ka, have been dated from offshore areas close to sjælland (bennike fig. 3 georadar profiles with interpretation below. the depth scale for profile 1 is in metres below sea level (m b.s.l.). profile 1 is located at gr shown in fig. 1, and profile 2 at 679 in fig. 1. https://doi.org/10.34194/geusb.v44.4630 http://www.geusbulletin.org bennike et al. 2020: geus bulletin 44. 4630. https://doi.org/10.34194/geusb.v44.4630 5 of 6 www.geusbul let in.org & jensen 1995; bennike unpublished data). the fossil assemblages indicate a tree-less open tundra-like vegetation with dwarf-shrub heaths. in the early holocene, lammefjorden was probably covered with lakes, bogs and forests. in the mid-holocene, the sea transgressed low-lying parts of odsherred, and lammefjorden was formed. the oldest age of marine shells from lammefjorden is c. 6.7 cal. ka, providing a minimum age for initial marine transgression. the age is at least 1000 years younger than expected when compared with other dated samples from the region (olsen et al. 2009; bennike et al. 2017). however, dating more samples might produce older ages. it is also possible that there is a shallow threshold to the inner part of lammefjorden, which prevented marine waters from flooding the area earlier. during the mid-holocene, marine faunas with large ostrea edulis indicate higher salinities and summer temperatures than at present and likely a larger tidal amplitude. in the mid-holocene, the rsl in the region was c. 3.5 m higher than at present, which would also lead to rich marine faunas. samples from several cores collected from deep basins in this study contained species-poor faunas. several explanations are offered but we consider high or very high sedimentation rates as the main cause. high sedimentation would have meant turbid water, making life difficult for benthic invertebrates. a shell bank dominated by mytilus and ostrea gave ages of 4.5 and 3.1 cal. ka. several hundred samples of ostrea shells from denmark have been dated and the ages show a peak of about 6.0 cal. ka (bennike et al. 2017). the common presence of the species in lammefjorden at a much later stage is surprising. the species immigrated to denmark in the early holocene and today it is absent from the inner danish waters (bennike et al. 2017). in supplementary file s5, ages of macrofossils versus depth are plotted. note that these samples were collected from a large area (c. 50 km2) with slightly different isostatic uplift rates. also plotted are global eustatic sea-level changes (lambeck et al. 2014) and a rsl curve for the marthe flak region (bennike unpublished data), located in an area with a similar marine limit to lammefjorden. as expected, the marine samples are plotted below the rsl curves. most of the terrestrial and lacustrine samples are located far above the curve, indicative of lakes and bogs that existed in the area long before the marine transgression. two samples (k-5308 and k-5996) appear to represent material that was formed just before the marine transgression. one sample of charcoal (k-5383) plots below the curve, indicating it was reworked. finally, k-3781 appears far below the curve. this sample, which comes from a 10 m long quercus stem, could have floated to the site and sank in deep water. conclusions late glacial and holocene lake and bog deposits occur below the marine deposits in the reclaimed lammefjorden. marine gyttja in some basin areas show potentially high sedimentation rates, which led to locally sparse benthic faunas. obtaining more radiocarbon dates would further constrain these sedimentation rates. in other areas, rich marine mollusc faunas were found, with large ostrea edulis, indicative of high summer temperatures, salinity and strong tidal currents. a marine shell at 6.7 cal. ka provides a minimum age for the marine transgression of lammefjorden. acknowledgements we thank two reviewers (jonathan lewis from loughborough university, uk and another anonymous reviewer) for their constructive and positive reviews of the manuscript. additional information funding statement this study was supported by geopark odsherred. author contributions ob: macrofossil analyses, manuscript writing. prj: field work and reporting on lithostratigraphy and georadar data. jwh: editing. additional files additional files are available online: https://doi.org/10.34194/geusb. v44.4630 references bennike, o. & jensen, j.b. 1995: near shore baltic ice lake deposits in faksebugt, southeast denmark. boreas 24, 185–195. https://doi. org/10.1111/j.1502-3885.1995.tb00772.x bennike, o. et al. 2004: lateand postglacial history of the great belt, denmark. boreas 33, 18–33. https://doi.org/10.1111/j.1502-3885.2004. tb00993.x bennike, o., pantmann, p. & aarsleff, e. 2017: holocene development of the arresø area, north-east sjælland, denmark. bulletin of the geological society of denmark 65, 25–35. christensen, c. 1994: lammefjorden. undersøgelser på fire lokaliteter ifjorden giver informationer om havniveauændringer og afkræfter formodet forekomst af tektoniske bevægelser af landjorden i atlantisk tid. nnu rapport 19994(16), 31 pp. christensen, c. & andreasen, e.a. 1999: strandforskydning i nordvestsjælland i atlantisk ogsubboreal tid. nnu rapport 19992(2), 19 pp. houmark-nielsen, m. et al. 2012: cosmogenic surface exposure dating the last deglaciation in denmark: discrepancies with independent age constraints suggest delayed periglacial landform stabilisation. quaternary geochronology 13, 1–17. https://doi. org/10.1016/j.quageo.2012.08.006 jakobsen, p.r. & bennike, o. 2017: geologisk model for lammefjordens inderlavning. danmarks og grønlands geologiske undersøgelse rapport 2017/39, 41 pp. jakobsen, p.r. et al. 2015: geopark odsherred. danmarks føste geopark. geoviden 2015(1), 20 pp. jakobsen, p.r. et al. 2016: boringer i lammefjordens inderlavning. danmarks og grønlands geologiske undersøgelse rapport 2016/41, 32 pp. https://doi.org/10.34194/geusb.v44.4630 http://www.geusbulletin.org https://doi.org/10.1111/j.1502-3885.1995.tb00772.x https://doi.org/10.1111/j.1502-3885.1995.tb00772.x https://doi.org/10.1111/j.1502-3885.2004.tb00993.x https://doi.org/10.1111/j.1502-3885.2004.tb00993.x https://doi.org/10.1016/j.quageo.2012.08.006 https://doi.org/10.1016/j.quageo.2012.08.006 bennike et al. 2020: geus bulletin 44. 4630. https://doi.org/10.34194/geusb.v44.4630 6 of 6 www.geusbul let in.org lambeck, k. et al. 2014: sea level and global ice volumes from the last glacial maximum to the holocene. proceedings of the national academy of the united states of america (pnas) 111, 15296–15303. http:// www.pnas.org/cgi/doi/10.1073/pnas.1411762111 madsen, a.p. et al. 1900: affaldsdynger fra stenalderen i danmark: undersøgte for nationalmuseet, 196 pp. copenhagen: c.a. reitzel. mertz, e.l. 1924: oversigt over de senog postglaciale niveauforandringer i danmark. danmarks geologiske undersøgelse ii. række, 41, 49 pp. olsen, j., rasmussen, p. & heinemeier, j. 2009: holocene temporal and spatial variation in the radiocarbon reservoir age of three danish fjords. boreas 38, 458–470. https://doi.org/10.1111/j.1502-3885. 2009.00088.x rørdam, k. & milthers, v. 1900: beskrivelse til geologisk kort over danmark. kortbladene sejrø, nykøbing, kalundborg og holbæk. danmarks geologiske undersøgelse i. række, 8, 132 pp. simonsen, p. 1946: stenalderbopladsen i gislinge lammefjord. historisk årbog for holbæk amt 1946, 38–61. troels-smith, j. 1942: geologisk datering af dyrholm-fundet. in: mathiassen, th., degerbøl, m. & troels-smidt, j. 1942: dyrholmen. en stenalderboplads på djursland. det kongelige danske videnskabernes selskab, arkæologisk-kunsthistoriske skrifter 1(1), 212 pp. westerby, e. 1933: nogle stenalderfund fra tørlagt havbund. meddelelser fra dansk geologisk forening 8, 231–248. https://doi.org/10.34194/geusb.v44.4630 http://www.geusbulletin.org http://www.pnas.org/cgi/doi/10.1073/pnas.1411762111 http://www.pnas.org/cgi/doi/10.1073/pnas.1411762111 https://doi.org/10.1111/j.1502-3885.2009.00088.x https://doi.org/10.1111/j.1502-3885.2009.00088.x geological survey of denmark and greenland bulletin 38, 2017, 57-60 57 camp century was a military base constructed by the us army corps of engineers (usace) in 1959 in the nearsurface layers of the greenland ice sheet at 77.13°n and 61.03°w and 1910 metres above sea level (clark 1965). the c. 55 ha base housed between 85 and 200 soldiers and was continuously occupied until 1964 (fig.1). camp century primarily served as an experimental facility for the usace to test ice-sheet construction concepts. recent danish scholarship has documented the political and military history of camp century in substantial detail (petersen 2007; nielsen & nielsen 2016). to summarise, project iceworm, the us army ambition to deploy offensive missiles within the ice sheet, was never realised. after three years of seasonal operation, camp century was finally abandoned with minimal decommissioning in 1967. the government of denmark has now established a geus-led programme for long-term climate monitoring, as well as one-time waste mapping, at camp century. here, we briefly review the historical scientific activities at camp century and introduce the future goals of the camp century climate monitoring programme. finally, we discuss the challenges and outlook of climate monitoring and waste mapping at the former military site. scientific heritage the usace conducted extensive glaciological and climatological research during the operation of camp century (fig. 2). much of their glaciological research focused on characterising the strength and density of the relatively porous, near-surface ice-sheet layer known as firn. these projects included measuring the deformational closure rates of near-surface tunnels (clark 1965), and excavating an inclined tunnel to 100 m depth to measure firn properties (kovacs et al. 1969). much of their climatological research focused on characterising spatial and temporal variability in snowfall. these projects include extensive surveys of regional snow-accumulation rates (mock 1968), and maintaining a continuous weather station record from october 1960 to august 1964 that remains unpublished. new programme for climate monitoring at camp century, greenland william colgan, signe b. andersen, dirk van as, jason e. box and søren gregersen 1960 10 20 30 40 50 60 70 liquid waste solid waste ice firn air 0 2015 2090 d ep th (m ) c • •thule airbasethule air base camp century 500 km 60°w 20°w 60°n 70°n a ••••••• 400 m 85 68 00 85 67 25 0 n b546750 547500 fig. 1. a: location of thule air base and camp century in north-west greenland. b: camp century as-built map with estimated georeferencing to 1960 (grey) and 2020 (black) locations in polar stereographic projection (epsg 3413). decadal borehole positions from 1960 to 2020 shown in red (colgan et al. 2016). blue lines denote a local coordinate system. c: estimated depths of solid and refrozen liquid wastes in the firn and ice beneath camp century in 1960, 2015 and 2090 (colgan et al. 2016). © 2017 geus. geological survey of denmark and greenland bulletin 38, 57–60. open access: www.geus.dk/publications/bull 5858 today, the usace-facilitated science at camp century is perhaps best known for producing the first systematic classification of ice-sheet snow facies (benson 1962), and recovering the first ice core to the bed of the greenland ice sheet (dansgaard et al. 1969). both the data and interpretations from these seminal studies continue to be highly cited today. after the closure of camp century, the us air national guard continued to use the camp century skiway, renaming it the greenland ice sheet training site (gits). aside from re-surveying the borehole position in 1977 and 1986 (gundestrup et al. 1987), there appears to have been virtually no data collected at camp century between the abandonment of the base in 1967 and the start of nasa program for arctic regional climate assessment (parca) activities at the site in 1993. parca activities included: deploying an automatic weather station in 1995 to record meteorology at the site (steffen & box 2001), drilling a 120 m deep ice core in 1996 to reconstruct snow accumulation rates (mosley-thompson et al. 2001), and measuring in-situ firn compaction rates in 1995–1996 (hamilton & whillans 2000). in 2010, a 35 m deep ice core was once again recovered at camp century, to further update snow accumulation and ice chemistry records since the termination of the usace ice-core record (buchardt et al. 2012). parca began regular airborne measurements of ice-surface elevation at camp century in 1993 (krabill et al. 2000), with nasa operation icebridge regularly collecting ice-penetrating radar data over the site since 2010 (leuschen et al. 2014). us national science foundation traverses from thule air base (ab) to summit station, which have approximately followed the usace trail to camp century since 2008, have been used as a science platform to measure accumulation rates (hawley et al. 2014). mapping and monitoring in 2016, geus participated in a multi-nation study that presented regional climate model simulations that suggested the ice-sheet surface mass balance at camp century may change from net snowfall to net melt by year 2100 under the un intergovernmental panel on climate change (ipcc) rcp8.5 ‘business-as-usual’ climate scenario (colgan et al. 2016). however, under the emissions mitigation characterised by the rcp4.5 climate scenario, net snowfall would persist at camp century until 2100. while colgan et al. (2016) also provided preliminary estimates of the non-trivial quantities of physical, chemical, biological and radiological wastes presently residing within the firn at camp century, at depths of between 35 and 65 m, perhaps the most socially significant outcome of this study was suggesting that the assumption that the abandoned base would be preserved for eternity by perpetually accumulating snowfall was no longer valid under the full range of ipcc climate pathways. in response to concerns from the government of greenland over the potential remobilisation of contaminants from camp century within the next century, the government of denmark has now established a programme for long-term climate monitoring, as well as one-time waste mapping, at camp century. this camp century climate monitoring programme will be led by geus and has four main goals: 1) to continuously monitor relevant climate variables, including the depth to which meltwater percolates, at the camp century site. this goal will be accomplished by installing an automated weather station that measures standard climatological variables controlling meltwater production (citterio et al. 2015). station measurements will be supplemented fig. 2. usace-applied glaciology research at camp century. left: measuring the deformation of firn in 1961. right: measuring the compressive strength of firn in 1964. photos: søren gregersen. 59 by thermistor strings to monitor deep firn temperatures, as well as observations of firn density and compaction profiles. 2) to regularly update annual likelihoods of meltwater interacting with abandoned materials at the camp century site over the next century. this goal will be accomplished by using a physically-based numerical model that couples meltwater percolation and firn evolution (charalampidis et al. 2016). this model will be forced by ipcc climate pathways and continuously improved using in situ observations, as well as novel parameterisations from community models. 3) to map the estimated spatial extent and vertical depth of abandoned wastes across the camp century site. this goal will be accomplished by using ice-penetrating radar and global positioning system measurements to map the camp century debris field during a one-time field campaign (machguth et al. 2016). delineating the present-day location of key infrastructure features will enable georeferencing of historical site maps. 4) to publicly report all findings from the camp century climate monitoring programme in a timely manner. this goal will be accomplished by streaming the data collected by sensors deployed at camp century in near-real-time, maintaining an internet outreach presence of the programme, and regularly publishing geus reports and papers in open-access, peer-reviewed journals. the camp century climate monitoring programme will undertake initial fieldwork at camp century during the summer of 2017, to deploy automated climate and firn sensors and collect ice-penetrating radar and firn-core observations (fig. 3). subsequent fieldwork at camp century will be undertaken, as needed, to service deployed instrumentation. during subsequent site visits, additional ice-penetrating radar data may potentially be collected in more concentrated areas of the debris field. the analysis of climate measurements, including firn temperatures, as well as numerical modelling of future meltwater percolation depths, will begin during the autumn of 2017, with anticipated first public reporting in the summer of 2018. near-real-time measurements from camp century, as well as programme outreach materials and publications, can be accessed at www.campcenturyclimate.dk. programme outlook geus has a long tradition of applied glaciology research. recent applied glaciology work includes operating the programme for monitoring of the greenland ice sheet (promice) on behalf of the government of denmark (ahlstrøm et al. 2008a), consulting for the government of greenland on the hydropower potentials associated with ice-sheet runoff (ahlstrøm et al. 2008b), and a growing involvement in private sector proglacial mining projects (citterio et al. 2009). with unique applied glaciology expertise gained through these and other operations, especially inhouse development of robust automated ice-sheet instrumentation and previous dedicated ice-coring and radar-acquisition campaigns, geus is well-suited to lead the camp century climate monitoring programme. indeed, geus involvement with camp century stretches from its operational period, when geus emeriti anker weidick and søren gregersen participated in research at the site, to contributing to the preliminary waste inventory and climate projections of the camp century site published last year. while the fundamental glaciology and climatology research performed by the usace gives camp century an unparalleled scientific heritage amongst greenland research fig. 3. left: servicing a promice automated weather station in the ice-sheet accumulation area in 2016. right: measuring firn density from a shallow borehole into which thermistors were installed in 2016. photos: baptiste vandecrux. 6060 sites, the military history of camp century gives the site unanticipated social significance in light of climate change. long-term climate monitoring, and one-time waste surveying, of camp century will provide danish and greenlandic stakeholders open access to relevant in-situ measurements and model projections. refined knowledge of the spatial and depth distribution of different wastes, as well as the changes in firn structure and meltwater production anticipated under climate change, will facilitate a science-based discussion of the shifting fate of camp century. at the broadest level, a better understanding of the implications of climate change on camp century will perhaps provide a better understanding of the importance of mitigating greenhouse-gas emissions, and averting, rather than adapting to the consequences of business-as-usual climate change. acknowledgements the camp century climate monitoring programme is funded by the danish ministry for energy, utilities and climate and by geus. references ahlstrøm, a. & the promice team 2008a: a new programme for monitoring the mass loss of the greenland ice sheet. geological survey of denmark and greenland bulletin 15, 61–64. ahlstrøm, a., mottram, r., nielsen, c., reeh, n. & andersen, s. 2008b: evaluation of the future hydropower potential at paakitsoq, ilulissat, w. greenland. danmarks og grønlands geologiske undersøgelse rapport 2008/37, 50 pp. benson, c. 1962: stratigraphic studies in the snow and firn of the greenland ice sheet. cold regions research and engineering laboratory. research report 70. buchardt, s., clausen, h., vinther, b. & dahl-jensen, d. 2012: investigating the past and recent delta18o-accumulation relationship seen in greenland ice cores. climate of the past 8, 2053–2059. charalampidis, c., van as, d., colgan, w.t., fausto, r.s., macferrin, m. & machguth, h. 2016: thermal tracing of retained meltwater in the lower accumulation area of the southwestern greenland ice sheet. annals of glaciology 57(72) 1–10. citterio, m., mottram, r., larsen, s. & ahlstrøm, a. 2009: glaciological investigations at the malmbjerg mining prospect, central east greenland. geological survey of denmark and greenland bulletin 17, 73–76. citterio, m., van as, d., ahlstrøm, a.p., andersen, m.l., andersen, s.b., box, j.e., charalampidis, c., colgan, w., fausto, r.s., nielsen, s. & veicherts, m. 2015: automatic weather stations for basic and applied glaciological research. geological survey of denmark and greenland bulletin 33, 69–72. clark, e.f. 1965: camp century evolution of concept and history of design, construction and performance. cold regions research and engineering laboratory. technical report 174, 69 pp. colgan, w., machguth, h., macferrin, m., colgan, j., van as, d. & macgregor, j. 2016: the abandoned ice sheet base at camp century, greenland, in a warming climate. geophysical research letters 43, 8091–8096. dansgaard, w., johnsen, s.j., møller, j. & langway, c.c. 1969: one thousand centuries of climatic record from camp century on the greenland ice sheet. science 166(3903), 377–380. gundestrup, n.s., clausen, h.b., hansen, b.l. & rand, j. 1987: camp century survey 1986. cold regions science and technology 14(3), 281–288. hamilton, g. & whillans, i. 2000: point measurements of mass balance of the greenland ice sheet using precision vertical global positioning system (gps) surveys. journal of geophysical research 105, 16,295–16,301. hawley, r.l., courville, z.r., kehrl, l.m., lutz, e.r., osterberg, e.c., overly, t.b. and wong, g.j. 2014: recent accumulation variability in northwest greenland from ground-penetrating radar and shallow cores along the greenland inland traverse. journal of glaciology 60(220), 375–382. kovacs, a., weeks, w.f. & michitti, f. 1969: variation of some chanical properties of polar snow, camp century, greenland. usa cold regions research and engineering laboratory, research report 276, 33 pp. krabill, w., abdalati, w., frederick, e., manizade, s., martin, c., sonntag, j., swift, r., thomas, r., wright, w. & yungel, j. 2000: greenland ice sheet: high-elevation balance and peripheral thinning. science 289, http://dx.doi.org/10.1126/science.289.5478.428 leuschen, c., gogineni, p., hale, r., paden, j., rodriguez, f., panzer, b. & gomez, d. 2014, updated 2016: icebridge mcords l1b geolocated radar echo strength profiles, version 2, [indicate subset used]. boulder, colorado usa: national snow and ice data center. http://dx.doi.org/10.5067/90s1xzrbax5n machguth, h., macferrin, m., van as, d., box, j., charalampidis, c., colgan, w., fausto, r., meijer, h., mosley-thompson, e. & van de wal, r. 2016: greenland meltwater storage in firn limited by nearsurface ice formation. nature climate change 6, 390–393. mock, s. 1968: snow accumulation studies on the thule peninsula, greenland. journal of glaciology 7, 59–76. mosley-thompson, e., mcconnell, j., bales, r., li, z., lin, p., steffen, k., thompson, l., edwards, r. & bathke, d. 2001: local to regional-scale variability of annual net accumulation on the greenland ice sheet from parca cores. journal of geophysical research 106, 33,839–33,851. nielsen, h. & nielsen, k. 2016: camp century – cold war city under the ice. in: doel, r., harper, k. & heymann, m. (eds): exploring greenland: cold war science and technology on ice, 195–216. palgrave studies in the history of science and tehcnology, palgrave macmillan us. petersen, n. 2007: the iceman that never came. ‘project iceworm’, the search for a nato deterrent, and denmark, 1960–1962. journal of scandinavian history 33, 75–98. steffen, k. & box, j. 2001: surface climatology of the greenland ice sheet: greenland climate network 1995–1999. journal of geophysical research 106, 33,951–33,964. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark; e-mail: wic@geus.dk geological survey of denmark and greenland bulletin 38, 2017, 41-44 41 numerous studies have proven that conceptual targeting based on integration of various geo-datasets can aid exploration companies to identify exploration targets (e.g. joly et al. 2013). this is particularly true in remote, underexplored areas that are commonly just covered by airborne geophysics and remote sensing and mapped geologically only on a regional scale. such regions are ‘exploration greenfields’ and may possess undiscovered economic deposits. the ice-free coastal strip of the archaean craton in south-east greenland overprinted by palaeoproterozoic orogeny (fig. 1) is such an area due to its remoteness and arctic-alpine conditions; and deep-seated, repeatedly reactivated structures and new magmatic episodes make large parts of this region potential for orogenic au occurrences. although only minor au mineralisation has been found to date, a large number of au-bearing rock samples (petersen & thomsen 2014) and stream sediment anomalies suggest an elevated potential particularly in the tasiilaq area (fig. 1b). a large field mapping campaign (kolb et al. 2016) and regional airborne magnetic surveys (riisager & rasmussen 2014) were conducted from 2012 to 2015, resulting in a uniform coverage of relevant geological and geophysical information, which can be combined with satellite remote sensing data. it was therefore decided to apply a fuzzy lo gic-based mineral prospectivity mapping (mpm) procedure (see below and fig. 2) combined with a mineral system approach of orogenic au (mccuaig et al. 2010). however, it was a challenge to apply this approach for several reasons: (1) strong topographical variations in this region distort some of the evidential datasets used in the mpm and their effect has to be minimised (e.g. aeromagnetic data are strongly affected by flight height). (2) fjords, ocean and glaciers strongly limit the accessible area for prospection, lead to a non-uniform data coverage for many datasets and affect the accuracy of evidential maps associated with the mpm. (3) lack of known in-situ au mineralisation limits the validation of final prospectivity maps. (4) the geological history is not fully understood, making it difficult to discard irrelevant features (e.g. magnetic anomalies may be related to insignificant palaeogene dykes or important older faults). such difficulties occur in many parts of greenland, and this study gives an idea of how meaningful mpm studies might be in other regions. regional geology the study area comprises the archaean north atlantic craton (nac) in the south and the palaeoproterozoic prospectivity mapping for orogenic gold in south-east greenland björn h. heincke and bo møller stensgaard geophysics aeromagnetic data a b geochemistry (stream sediments) au concentrations anomalies associated with 2nd faults tasiilaq anomalies associated with tertiary dykes trends from regional magnetic compilation camp-m 200 km au [ppb] 1–4 4–10 10–20 20–44 44–190 greenland fig. 1. examples of data types used for the mpm study in south-east greenland (framed area in index map). a: mapped elongate anomalies from aeromagnetic data (see background map in riisager & rasmussen 2014) and regional magnetic compilations used as proxies for faults of 1st and 2nd order in the critical processes ‘source’ and ‘pathways’. b: stream sediment au concentrations used as a proxy for the ‘chemical scrubber’. black lines in a and b mark the area used in the interference network. © 2017 geus. geological survey of denmark and greenland bulletin 38, 41–44. open access: www.geus.dk/publications/bull 4242 nagssugtoqidian orogen in the north (figs 1, 3; kolb et al. 2016). to the north, the orogen includes an archaean foreland that comprises rocks from the rae craton. to the south, the margin of the nac is affected by deformation and intrusives of the palaeoproterozoic ketilidian orogen. the nagssugtoqidian orogen comprises tectonically reworked archaean rocks subjected to high-grade metamorphism and slivers and belts of palaeoproterozoic metavolcanic, metasedimentary and intrusive rocks. on the basis of differences in lithologies and tectono-metamorphic history, the orogen is divided into four terranes, from north to south the isortoq terrane, the ammassalik intrusive complex (aic) and the kuummiut and schweizerland terranes. the nac is dominated by felsic orthogneisses with subordinate supracrustal rocks, with synto post-tectonic alkaline intrusions in the skjoldungen area (63°10´–63°40´n). main deformation events within the nac are the timmiarmiut and skjold-ungen orogenies. continental breakup in the palaeogene led to the emplacement of coast-parallel dykes in the northern area (north of 64°n) that can clearly be identified as anomalies in aeromagnetic data (red lines in fig. 1a). the mineral prospectivity mapping approach mineral prospectivity maps used for targeted exploration highlight areas with coincident geological features that are important for a given commodity. for our mineral prospectivity mapping we use geoscience data that were selected on the basis of a mineral system approach for orogenic gold. the mineral system approach provides a holistic view of the critical geological, physical and chemical processes needed to generate a mineral deposit. for orogenic au systems mccuaig et al. (2010) identified the following critical processes: (i) a source of au in the upper mantle, (ii) active pathways allowing fluids to flow through the crust, (iii) physical traps in which fluids are throttled and focused and (iv) a chemical ‘scrubber’ associated with hydrothermal mineralisation and alteration (fig. 2). since ore-forming processes typically cannot be directly mapped, they must be inferred from geological features referred to as targeting elements. these are rarely directly measurable but are approximated from responses in geoscience data such as faults estimated from magnetic anomalies. these proxies are represented in this approach by uniform spatial grids named evidential maps (em) that can easily be combined to build prospectivity maps. to quantify and combine all information we use ‘fuzzy logic’ operations which have proven in many cases (e.g. joly et al. 2013) to be suitable for building meaningful mineral prospectivity maps. first, all ems were rescaled (i.e. transformed into maps with values ranging from 0 to 1) by employing so-called fuzzy membership functions to obtain representative maps describing how far the prospectivity is supported (0.0 = not at all; 1.0 = fully). in a multistage interference network, fuzzy or (maximum operator equivalent to logical union) and and (minimum operator equivalent to logical intersection) operators were then applied to combine the ems to create first intermediate fuzzified maps representing the four critical processes and finally fuzzy prospectivity maps used to predict areas of high au potential (electronic supplementary (es) figure: fig. es1). further details about the mineral prospecitivty mapping procedure, the underlying mineral system approach and data sets are given in stensgaard & heincke (2016). calculation of prospectivity maps most of the ems associated with structural information were derived from anomalies in potential field data, either from regional data compilations (gaina et al. 2011) or recent aeromagnetic surveys (fig. 1a). a 1:500 000 scale digital geological map (stensgaard et al. 2016) was used critical processes constituent processes targeting elements proxies (evidential maps) source pathways physical trap/ throttle chemical trap/ scrubber tapping of mantle source magmatic input to crust faults/shear zones pathways along unit interfaces enhanced permeability/focus hydrothermal fluid/mineralisation hydrothermal alteration deep-seated structures intrusions (as access to source areas) adjacent faults/ shear zones i) 1st order ii) 2nd/3rd order ’interface complexity structural intersection ’structural complexity’ geochemical signature favourable for or indicative of mineralisation chemically favourable host rock for au mineralisation lineaments from regional a) gravity b) magnetic mapped intrusions lineaments from airborne magnetic data density of mapped unit interfaces intersections of lineaments (from airborne magnetic) intersection of lineaments with mapped geological units density of geological units stream sediment geochemistry (au, as, w, mo, sb cs, bi, ratio of ni/ mg) mapped rock units a) mafic units b) supracrustal units c) marble units satellite aster data a) mafic index b) carbonite index c) silica index regional scale district scale regional/ district scale prospect scale regional scale district scale regional scale prospect scale fig. 2. the mineral system model used for the orogenic gold deposit type. the critical processes are associated with a series of constituent processes, targeting elements and proxies. 43 to extract additional structural information (fig. es2a) and to identify geological units and other settings favourable for au mineralisation, e.g. intersections between units and cross-cutting structural elements. other ems associated with preferable rock types were obtained from mineral indices of aster satellite data (fig. es2b). finally, the contents of au and geochemical pathfinder elements for au mineralisation (as, sb, cs, w, bi, mo, ni/mg) in stream-sediment samples (fig. 1b) were presented in ems that reflect relevant hydrothermal mineralising fluids as well as associated alteration halos. the non-uniform spatial data coverage made it necessary to use interpolation (kriging and natural neighbour gridding) to create the related ems. to rescale the ems we used fuzzy membership functions that were estimated on the basis of qualitative and quantitative knowledge (see fig. es3). the region used to determine the ems and perform the interference network calculations comprises both the ice-free onshore, adjacent offshore and ice-covered areas; however, only accessible ice-free areas were considered in the evaluation of prospectivity maps (figs 3a, b). there are uncertainties in all steps of the mineral prospecitivity mapping procedure (mccuaig et al. 2010), and it is important to test how reliable the final prospectivity maps are. if known mineral occurrences that originate from the assumed mineral system are present, one option is to construct prediction-rate curves to evaluate and adapt the prospectivity procedure (carranza 2009). however, in the absence of such mineral occurrences, we used streamsediment locations with au concentrations > 20 ppb as ‘deposits’ in such curves (see circles in figs 3a, c). this led to a number of inaccuracies, as sample locations typically do not coincide with locations in the catchment area where gold was eroded, and where no unique and simple function generally links concentrations of au occurrences and stream-sediment samples. hence, conclusions based on such validation should be considered carefully. to evaluate the robustness of the mineral prospectivity mapping and the relevance of different targeting elements, gabbro marble diorite, tonalite brown orthogneiss ttg gneiss ultramafic rocks ultramafic and metasedimentary rocks meta-anorthosite/ leucogabbro paragneiss, amphibolite, ultramafic rocks contact-metamorphic anatectic gneiss norite, gabbro, diorite, granodiorite granite, granodiorite, diorite palaeoproterozoic archaean 38°00 66°30 66°n 66°30 66° 65°30’ 37°00 36°00 38° 37° 36°w tasiilaq kuummiut terrane schweizerland terrane ammassalik intrusive complex kap tycho brahe se rm ilik 25 km 0 20 40 60 80 100 0 20 40 60 80 100 cumulative proportion of area with decreasing score fuzzy value high low c um ul at ive p ro po rt io n of st re am se di m en t s am pl es w ith > 2 0 pp b au (i n %) prediction−rate curve 67°n 66°n 65°n 64°n 43°w 39°w 35°wa b d c i ii 100 km fig. 3. a: final prospectivity map obtained by using all evidential maps. white circles: stream sediment locations with au > 20 ppb. red circle: the sample with highest concentration of 196 ppb. b: validation curves showing the cumulative area with decreasing score versus the cumulative number of stream sediment samples with > 20 ppb au in the same area. (in the blue and the green curves, ems associated with streamsediment data from au and all elements are not incorporated). c: map of the area around tasiilaq (red rectangle in a). pixels with a high fuzzy value of > 0.85 are shown in red. d: geological map of the same area. 4444 we repeated the procedure with several different combinations of ems (e.g. all ems, all ems except for the one based on au concentration from stream sediments). irrespective of the combination of ems, the resulting validation curves show that high fuzzy scores correlate strongly with elevated au in stream sediment samples (fig. 3b) suggesting that the mineral prospectivity mapping is reliable and robust. results and discussion to identify areas with high orogenic au potential, we highlighted pixels (pixel size: 3 × 3 km) in the prospectivity map with high fuzzy values > 0.85 (red colour in fig. 3c) associated with ~8% of the total area. particularly the tasiilaq (i) and the skjoldungen (ii) regions are characterised by high scores (fig. 3a). in region i, a couple of areas stand out as being anomalous in gold and its pathfinder elements; and locations of both stream sediment samples with high au concentrations (fig. 1b) and au-bearing rock samples (petersen & thomsen 2014) coincide well with areas having elevated membership fuzzy values. the highest scores are aligned along a wnw–ese-oriented corridor that coincides with the ammassalik intrusive complex (aic). in particular the boundaries of the aic to the kummuit and the isortoq terranes in the north and south are identified as prospective (fig. 3d; kolb 2014). a major structure that likely represents a suture zone which may have acted as a mantle source-tapping feature is reported along the northern boundary of the aic (kolb 2014) and is confirmed by geophysics (riisager & rasmussen 2014). other deepseated structures have been suggested south of the aic in the isortoq terrane (nutman et al. 2008). this means that an elevated au potential from the mineral prospectivity mapping is supported by the main geological settings. in contrast, no elevated au concentrations from samples or other indications for au occurrences exist within the cratonic skjoldungen region (ii). the latter area is characterised by predominant mafic granulites with thin belts of deep-crustal-formed mafic and ultramafic rocks considered unfavourable for orogenic au mineralisation (kolb et al. 2016). at this stage, the mineral system approach is set up to not exclude rock units formed at large depths, and it should be adjusted in future to take this information into account. in summary, the results of this study suggest that modern mineral prospectivity mapping schemes based on an mineral system approach can improve the geological understanding of mineral prospectivity even in regions that are not densely covered by all data types and underexplored. acknowledgements the project was jointly financed project by geus and the ministry of mineral resources, government of greenland. we thank s. weatherley and e.v. sørensen, geus, for comments and contributions. references carranza, e.j.m. 2008: geochemical anomaly and mineral prospectivity mapping in gis. handbook of exploration and environmental geochemistry 11, 351 pp. gaina, c., werner, s.c., saltus, r., maus, s. & the camp-gm group 2011: circum-arctic mapping project: new magnetic and gravity anomaly maps of the arctic. geological society (london) memoirs 35, 39–48. joly, a., dentith, m.c., porwal, a., spaggiari, c.v., tyler, i.m. & mccuaig, t.c. 2013: an integrated geological and geophysical study of the west arunta orogen and its mineral prospectivity. geological survey of western australia report 113, 89 pp. kolb, j. 2014: structure of the palaeoproterozoic nagssugtoqidian orogen, south-east greenland: model for the tectonic evolution. precambrian research 255, 809–822. kolb, j., stensgaard, b.m. & kokfelt, t.f. 2016 (eds): geology and mineral potential of south-east greenland, danmarks og grønlands geologiske undersøgelse rapport 2016/38, 157 pp. mccuaig, t.c., beresford, s. & hronky, j. 2010: translating the mineral systems approach into an effective exploration targeting system. ore geology reviews 38, 128–138. nutman, a.p., kalsbeek, f. & friend, c.r.l. 2008: the nagssugtoqidian orogen in south-east greenland: evidence for paleoproterozoic collision and plate assembly. american journal of science 308, 529–572. petersen, j. & thomsen, l.l. 2014: results in south-east greenland from ujarassiorit-program. danmarks og grønlands geologiske undersøgelse rapport 2014/79, 65–66. riisager, p. & rasmussen, t.m. 2014: aeromagnetic survey in southeastern greenland: project aeromag 2013. geological survey of denmark and greenland bulletin 31, 63–66. stensgaard, b.m., kolb, j., kokfelt, t.f. & klausen, m.b. 2016: digital revised 1:500 000 geologic map of south-east greenland 62°00´n to 67°00´n and 33°00 w to 44°00 w. copenhagen: geological survey of denmark and greenland. stensgaard, b.m. & heincke, b.h. 2016: targeting: prospectivity mapping for orogenic gold in south-east greenland. danmarks og grønlands geologiske undersøgelse rapport 2016/43, 176 pp. authors’ addresses b.h.h., geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350 copenhagen k, denmark, e-mail: bhm@geus.dk b.m.s., eit rawmaterials gmbh, europa center, tauentzienstr. 11, 10789 berlin, germany. local office: goldschmidtsvej 23, dk-2000, frederiksberg, denmark. geological survey of denmark and greenland bulletin 33, 2015, 1-7 1 geological survey of denmark and greenland bulletin 33 • 2015 review of survey activities 2014 edited by ole bennike, adam a. garde and w. stuart watt geological survey of denmark and greenland ministry of climate, energy and building 22 geological survey of denmark and greenland bulletin 33 keywords geological survey of denmark and greenland, survey organisations, current research, denmark, greenland. cover photographs from left to right 1 fractured granite on the island of bornholm. photograph: merete binderup. 2 th e swedish ice-breaker oden in the arctic ocean. photograph: martin jakobssen. 3 a geologist enjoying the magnifi cent view over the landscape in south-east greenland. photograph: jakob lautrup. 4 a geologist crossing a local ice cap on nuussuaq peninsula in west greenland. photograph: jason briner. frontispiece: facing page ice conditions in the area north of greenland during the geus lomrog 2012 cruise with the swedish ice-breaker oden. see paper by c. marcussen et al. in this issue. photograph: th omas funck. chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, department of geoscience, aarhus university; minik rosing, geological museum, university of copenhagen; finn surlyk, department of geosciences and natural resource management, university of copenhagen scientifi c editors: ole bennike, adam a. garde and w. stuart watt editorial secretary: jane holst referees (numbers refer to fi rst page of reviewed article): anonymous (17, 33); leon bagas, au (49); liz bagshaw, uk (69); niels balling, dk (21); sean brennan, usa (85); hilmar bungum, no (21); david burgess, ca (61); valeria caironi, it (53); mikael calner, se (9); jakob qvortrup christensen, dk (33); ida lykke fabricius, dk (13); xavier fettweis, be (65); peter gerling, de (85); henning haack, dk (25); maths halstensen, no (13); douglas r. hardy, usa (69); jens havskov, no (25); martin heinesen, fo (41); julie hollis, gl (53); niels tvis knudsen, dk (57); john korstgård, dk (17, 37); jeppe malmros, ch (57); ruth mottram, dk (65); asger ken pedersen, dk (45); minik rosing, dk (49); mark simoni, no (37); henrik stendal, gl (77, 81); svend stouge, dk (9); max strunck, se (73); phil symonds, au (41); leif th eilgaard, dk (29); henrik vedel, dk (29); tod e. waight, dk (45); jacob clement yde, no (61, 73). illustrations: benny m. schark, jette halskov, stefan sølberg, willy l. weng, susanne rømer and frants v. platen-hallermund layout and graphic production: kristian a. rasmussen printer: rosendahls-schultz grafi sk a/s, albertslund, denmark manuscripts received: 8 december 2014 – 12 may 2015 final versions approved: 27 january 2015 – 21 may 2015 printed: 7 july 2015 issn (print) 1604-8156 issn (online) 1904-4666 isbn (print) 978-87-7871404-6 isbn (online) 978-87-8771-405-3 citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 33, 88 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk and at www.geus.dk/publications/bull © de nationale geologiske undersøgelser for danmark og grønland (geus), 2015 for the full text of the geus copyright clause, please refer to www.geus.dk/publications/bull 3 33 relationship between groundwater chemistry and the precambrian basement rocks on eastern bornholm, denmark p. gravesen, r. jakobsen and b. nilsson 37 assessment of the mineral raw material potential in denmark – methods and major fi ndings j.k. keiding, p. kalvig, c. ditlefsen, s. lomholt and p.r. jakobsen 41 th e continental shelf project of the kingdom of denmark – status and issues c. marcussen, f. mørk, t. funck, w.l. weng and m. pedersen 45 magma mixing, mingling and hybridisation at diff erent crustal levels: snapshots from 1.9 billion years of magmatism in south-eastern greenland t.f. kokfelt, s.m. weatherley, j.k. keiding and t.b. árting 44 ghana greenland canada kenya burundi ethiopia nigeria cameroon malawi zambia bolivia bolivia denmark ghana brazil 7 review of survey activities 2014 flemming g. christiansen 9 th e lower palaeozoic now fully cored and logged on bornholm, denmark n.h. schovsbo, a.t. nielsen and k. klitten 13 down-hole permeability prediction – a chemometric wire-line log feasibility study from a north sea chalk well k.h. esbensen, n.h. schovsbo and l. kristensen 17 th rust-fault architecture of glaciotectonic complexes in denmark s.a.s. pedersen and l.o. boldreel 21 consistency of postglacial geodynamics for the kattegat region, southern scandinavia, based on seismological, geological and geodetic data s. gregersen and p.h. voss 25 acoustic events on a small seismological network – shock waves from thunder and fi reballs p.h. voss, t. dahl-jensen and t.b. larsen 29 a hydrological early warning system for denmark based on the national model h.j. henriksen, s. stisen, x. he and m.b. wiese grey indicates countries where geus had projects in 2014 orange indicates countries with geus projects decribed in this volume. 49 a quartz-wolframite-molybdenite vein and scheelite in amphibolite horizons from th rudvang peninsula, skjoldungen, se greenland d. rosa and t. ulrich 53 follow-up on ujarassiorit mineral hunt fi nds and outreach activities, south-east greenland m.d. poulsen, h. paulick, d. rosa, v.j. van hinsberg, j. petersen and l.l. th omsen 57 greenland ice sheet melt area from modis (2000– 2014) r.s. fausto, d. van as, j.a. antoft , j.e. box, w. colgan and the promice project team 61 unique applied glaciology challenges of proglacial mining w. colgan, h.h. th omsen and m. citterio 65 observed melt-season snowpack evolution on the greenland ice sheet c. charalampidis and d. van as 69 automatic weather stations for basic and applied glaciological research m. citterio, d.v. as, a.p. ahlstrøm, m.l. langer, s.b. andersen, j. e. box, c. charalampidis, w.t. colgan, r.s. fausto, s. nielsen and m. veicherts 73 digital models based on images taken with handheld cameras – examples on land, from the sea and on ice e.v. sørensen, m. bjerager and m. citterio 77 investigations of detrital zircon, rutile and titanite from present-day labrador drainage basins: fi ngerprinting the grenvillean front t.b. th omsen, c. knudsen and a.m. hinchey 81 composition of ilmenite and provenance of zircon in northern brazil c. knudsen, t.b. th omsen, f. kalsbeek, j.a. kristensen, h. vital and r.k. mclimans 85 reserves and resources for co2 storage in europe: the co2stop project n. poulsen, a. bocin-dumitriu, s. holloway, k. kirk, f. neele and n. smith 5 thailand vietnam phillippines indonesia singapore india tajikistan 7© 2015 geus. geological survey of denmark and greenland bulletin 33, 7–8. open access: www.geus.dk/publications/bull review of survey activities 2014 flemming g. christiansen deputy director over the past decades the geological survey of denmark and greenland (geus) has gradually changed from an organisation that focussed on basic mapping and research to one that provides knowledge for discussions and decisions within society and the political system in denmark and greenland. knowing the importance – and in some cases controversy – of many of these decisions, it is of the utmost importance that such input is based on high-quality and well-documented research that is objective, transparent and easily accessible. such input to important public and political discussion is also refl ected in this issue of the review of survey activities that contains a total of 20 four-page papers, eight on denmark, eighth on greenland and four on broader themes. in addition to many classical research summary papers, there are also examples of activities giving input to important political decisions on energy such as shale gas and carbon capture and storage technology, use of and exploration for mineral resources in denmark and greenland, climate (monitoring of the greenland ice sheet) and foreign policy (international boundaries). activities in denmark th e activities and research in denmark of geus cover a wide range of topics within our main programme areas: data, water, energy, mineral resources as well as nature and climate. th e island of bornholm has a diff erent geological history from the rest of denmark; a dominance of outcropping basement rocks and palaeozoic sedimentary rocks gives special research possibilities. one paper summarises information on the lower palaeozoic from scientifi c wells. th is succession is now fully cored and logged, thereby providing important data for correlation to other regions for assessing resources in denmark and scandinavia and for local use of water supplies. chalk is a very important rock type for denmark as it hosts more than 90% of the danish petroleum reserves in the north sea. one paper discusses one of the main challenges, namely prediction of permeability, which is a very critical parameter for production performance. geus is involved in many studies of quaternary and recent geological processes. one paper describes thrust-fault architecture of glacio-tectonic complexes using 3d geological models based on integrated photo-data from cliff sections with high resolution seismic data. geus records seismological events at many locations in denmark and greenland, data that are useful for many different purposes. one paper applies seismological, geological and geodetic data to discuss the consistency of postglacial geodynamics (especially uplift ) in the kattegat region. another paper provides examples of shock waves from thunder and fi reballs that have been detected on seismograms. th e use of groundwater is very important for the danish society, and geus carries out many studies on water resources and possible future scenarios due to changes in climate and use. one paper uses the comprehensive and constantly updated dk-model in the process of developing a hydrological early-warning system for denmark that can be important for water plans for emergency managers. another paper describes the relationship between groundwater chemistry and the weathering of precambrian basement rocks on eastern bornholm. th e center for minerals and materials (mima) at geus was established to identify and study important raw material chains from source to use, with the goal of enhancing our knowledge of the risk of resource scarcity and the ensuing vulnerability of the danish society. one paper gives an assessment of the mineral resource potential of the onand off shore areas in denmark with focus on methods and the main results. denmark has a high potential for aggregates, various clays, chalk and lime, salt and granite – but there are also many confl icting interests on the use of land and seabed where detailed planning and regulations are required. activities in greenland once again there was a high level of activity in and concerning greenland in 2014. many large and small projects were carried out, studies that are important for evaluating and marketing the resource potential in greenland. th e level of activity in both oil and mineral exploration is rather low at present, but it is very important to prepare for a future when 88 prices of the most important commodities will rise again. monitoring of ice and predicting climate changes are also important geus activities. th e work on the continental shelf project reached a new culmination point, when the 5th submission concerning a very large area north of greenland, with great international media coverage, was sent jointly to the commission on the limits of the continental shelf by the government of the kingdom of denmark and the government of greenland in december 2014. one paper gives a status for the recent work on the project and mentions some of the issues for the coming years where data and results will be presented and used in negotiations, maintainance and hopefully supplemented by additional scientifi c data. several papers focus on mapping and evaluation of the mineral potential in greenland. one paper gives snapshots of the magmatic history in south-eastern greenland with examples of magma mixing, mingling and hybridisation at diff erent crustal levels. a second paper gives details of mineralisation in the skjoldungen area, south-east greenland, where tungstenand molybdenum-bearing minerals have been discovered within veins in amphibolites. a third paper follows up on the annual ujarassiorit public mineral hunt in south-east greenland where a number of interesting samples have been found by local stone collectors (e.g. corundum, precious and base metals), and where considerable time during fi eld work was used for outreach activities in several small settlements. studies and monitoring of the greenland ice sheet and local glaciers provide a signifi cant contribution to models for global sea-level rise. th e important monitoring programme of the greenland ice sheet (promice) that was initiated in 2007 continuously supplies crucial data that are used in a number of subsequent projects and in key publications. th ree papers in this issue use such data from promice. one discusses the ice-sheet melt area, where data from modis (nasa’s moderate-resolution imaging spectroradiometer) are validated against promice data. a second is on the observed melt-season snowpack evolution of the greenland ice sheet. a third paper gives a technical description of the automatic weather stations that have been developed for basic and applied research. a fourth paper addresses applied glaciological challenges of proglacial mining, which may be important in greenland where several potential mining sites are located very close to the margin of the ice sheet. broader thematic activities internationally geus also works in many diff erent countries with a variety of projects and is involved in broader thematic studies. th e fi rst of these thematic papers describes how digital outcrop models can be made based on images from a handheld camera, with examples from land, sea and ice. a second paper gives details of a so-called provenance study in labrador, canada, using detailed analytical data on zircon, rutile and titanite that can be applied to map out boundaries between major orogenic terrains. another provenance study of heavy minerals from northern brazil is also presented, in this case with focus on titanium placer deposits. th e challenge of climate change demands reduction in global co2 emissions. one of the most promising technical solutions is to use carbon dioxide capture and storage (ccs). th e fi nal paper is a summary of the co2stop project, which has designed a database on storage capacity in 27 european countries and identifi ed important gaps in our knowledge. geological survey of denmark and greenland bulletin 28, 2013, 41-44 41 seismic activity in denmark: detection level and recent felt earthquakes trine dahl-jensen, peter h. voss, tine b. larsen and søren gregersen the geological survey of denmark and greenland (geus) records seismological data at six locations in denmark (fig. 1) and all data from these stations are manually reviewed for events like earthquakes and explosions. the identified events are analysed and located, in many cases using supporting data from stations outside denmark. seismic events have been recorded instrumentally in denmark since 1929, but earthquakes felt in denmark have been reported as far back as 1515 (lehmann 1956; gregersen et al. 1998; geus 2012). this article reports on the developments in detection level of both man-made events and natural earthquakes within the danish exclusive economic zone (eez) from 2000 to 2012. changes in detection level are mainly due to the availability of data from new seismic stations in sweden and norway as well as from a geus test station at gøttrup in nw jylland. as a case study, the list of events on and around bornholm is reviewed. also described here are the reported intensities at two recent felt events in denmark (north sea magnitude 4.3 on 19 february 2010 and kattegat magnitude 4.1 on 6 august 2012). development in detection level and completeness from 2000 to 2012 the events from 2000 to 2012 located within the danish eez (fig. 1) are divided into known explosions (pink), events recorded only on one or two stations which are typically small events (yellow), events where a magnitude could not be calculated (grey) and events which possibly are earthquakes (red) and recorded on three or more stations. many of the ‘possibly earthquakes’ events occur in areas where explosions are known to take place, and many of these events are suspected to have a man-made origin. many explosions are reported to the seismological service at geus, and are tagged as such, but many more are not reported and only some are tagged as possible explosions. the north sea is the only area with no explosions known to the seismological service but with many probable earthquakes. for example, the earthquake felt in february 2010, described below, is located in this area. however, earthquakes do occur in other areas of denmark; for example the felt earthquake that occurred in august 2012 in kattegat. fig. 1. seismic events within the danish exclusive economic zone (eez). among the recorded events are un-identified explosions, particularly in the area where many identified explosions are marked (in pink). the two felt events described in this paper are marked in blue. k = kattegat on 6 august 2012 and n = north sea on 19 february 2010. seismological stations are green triangles. the events are divided into events seen on three or more stations with defined magnitude (red), events seen on only one or two stations with defined magnitude (yellow), events where no magnitude has been calculated (grey) and known (or probable) explosions (pink). inset: the eez around bornholm after revision of the database. © 2013 geus. geological survey of denmark and greenland bulletin 28, 41–44. open access: www.geus.dk/publications/bull magnitude 1 2 3 4 station known explosions 1 or 2 stations no magnitude bsdbsd snart homb mud goet n k lldlld gidgid copcop bsdbsd 15˚e10˚e5˚e 59˚n 57˚n 55˚n 55˚n 15˚e 4242 in the years 2000–2005, fewer than 15 events (either earthquakes or non-reported explosions) observed on three or more stations were recorded each year, with an additional c. 20 known explosions. only large events were seen outside the danish network of stations (at the time bsd, cop, mud, lld and gid) (fig. 1). the number of events rose dramatically in 2006 (fig. 2), as a result of the installation of the norwegian seismological station snart (nnsn 2012). the addition of this station has aided in locating events, as it provides a much improved geometry of the station network (fig. 1). the increase in the number of events recorded is also due to a change in policy in 2006; since then events located by azimuthal analysis when only one or two stations have recorded the events (fig. 2) are included. for events seen on three or more stations, the increase is most pronounced for the smaller events, under magnitude 2, but also events with magnitudes between 2 and 3 are more numerous. the next large step up in event detections occurred in 2010 (fig. 2). this is due to data from the large swedish network (snsn 2012) becoming available, and also data from the new norwegian station homb. many of the additional events, only recorded on the snsn stations, have no magnitude due to missing calibration information from the new snsn. in 2012 the number of events with no magnitude declined, while the total number of events is fairly constant, as a result of snsn becoming established and complete metadata becoming available. in denmark we added goet as a test station in 2012, and data came online in november 2012. together with mud and the norwegian stations the azimuth coverage for the many events in the danish north sea is highly improved. explosions in the database far from all activity recorded within the danish eez are natural earthquakes. the seismological service at geus is frequently notified by the danish navy of upcoming or recent blasts. following world war ii, numerous unexploded mines and ammunition are still present in danish water; the largest neutralised in 2012 was equivalent to 800 kg tnt. the navy searches for the mines and detonates them on site. many are seen as signals on the seismic stations, and if known to be explosions they are logged as such in the database. on fig. 1 the known explosions are marked with pink, and are present in large parts of danish waters. known explosions also occur on land – for example in controlled-source scientific projects (thybo et al. 2006) and on rare occasions a house demolition. however, not all explosions are known by the seismological service. natural earthquakes are distributed evenly throughout the 24 hours of the day, while manmade events such as explosions mainly take place during the daytime. the navy usually blasts in the early afternoon. figure 3 illustrates the distribution of recorded events sorted by of day, and it is clear that the distribution is heavily skewed towards events in daytime hours. the known explosions are, as expected, concentrated during daytime hours; the exception being urgent blasts when a find of undetonated explosives endangers the surroundings. scientific blasts often take place during the quiet night hours. in the period 2000–2005 the events are evenly distributed through all 4 hour intervals, while the events in both the 2006–2009 and the 2010–2012 periods have a large overrepresentation in daytime. in total fig. 3. seismic events seen on three or more stations sorted by time of day. the events are sorted into 4 hour intervals in utc time. denmark is one hour ahead of utc (two hours in summertime). for each series of years, the number of events within a 4 hour interval is calculated as per year, so the three periods (2000–2005; 2006–2009 and 2010–2012) can be compared. in addition the number of known explosions in the entire period (2000–2012) is sorted in the same manner for comparison. inset: events in the area in the danish north sea with many events recorded but no known explosions. fig. 2. statistics on event detection from 2000–2012 within the danish exclusive economic zone. known or probable explosions are not included. 100 80 60 40 20 0 2000 2001 2002 2003 2004 2005 2006 2007 2008 2009 2010 2011 2012 1–2 stations magnitude > 4 magnitude 3–4 magnitude 2–3 magnitude 1–2 magnitude <1 no magnitude no. of events 0–4 4–8 8–12 12–16 16–20 20–24 0 10 20 30 40 known explosions 2000–2012 2000–2005 2006–2009 2010–2012 0 1 0–4 4–8 8–12 12–16 16–20 20–24 north sea 2000–2012 no. of events per year 43 426 events are included. assuming that the night-time level of events is correct, and the natural earthquakes are evenly distributed, a simple calculation shows that at least half the events are probably not natural earthquakes. not all explosions are reported to geus, for example explosions carried out by foreign naval vessels participating in exercises in danish waters or mines or ammunition neutralised by our neighbouring countries but erroneously located into danish waters. the only area where no known explosions are located is within the group of events in the danish north sea. for events in this area there is no concentration in daytime (fig. 3 – inset), and they are assumed to be natural earthquakes. case study bornholm a revision of the entire database of seismic events is in its initial phase, and for 2000 to 2012 the revision has been carried out on bornholm and within the eez around the island. here quarry blasts add to the man-made events recorded, as bornholm granit blasts several times a week at set times (paul ebbesen, rønne granit, ncc, pers. communication). by logging all events known or strongly suspected to be explosions, excluding small events seen only on one or two stations (and with a large uncertainty in location) and events so small that no magnitude could be calculated in spite of all necessary station information, only two events remain as probable natural earthquakes. the original list contained 129 events within the eez around bornholm, of which 25 were seen on three or more stations (of these three without magnitude). the remaining 104 events were seen only on one or two stations, and many with no magnitude. seventy events are suspected quarry blasts. in all 23 events had a magnitude of 2 or higher. ten previously known explosions make an original total of 139 events recorded. the two ‘surviving’ events are marked on the inset on fig. 1 and are a magnitude 2.0 event in 2006 and a very small magnitude 0.8 event in 2011. recent felt earthquakes it is rare for denmark to experience an earthquake which can be felt. however, it does happen and the two most recent felt earthquakes are briefly described below. north sea, 19 february 2010, magnitude 4.3 at 21:08 utc (22:08 local time) on 19 february 2010, a magnitude 4.3 earthquake occurred 45 km offshore the north-western danish coast at a depth of 39 km (fig. 4a). the focal mechanism indicates a reverse fault overthrusting to the west, in agreement with earlier earthquakes in the area. this is the most active area in denmark, with many known earthquakes. from the instrumental era, table 1 lists fig. 4. reports of observed intensity from the public. the earthquakes are marked by blue stars scaled by size. a: north sea on 19 february 2010 magnitude 4.3. geus received 344 reports from people who felt the earthquake. seven small aftershocks (magnitude 1.9 to 2.8) were observed during four weeks after the main event. b: kattegat on 6 august 2012 magnitude 4.1. in addition to reports received by geus (441), we have reports from the united states geological survey (in all 76 reports of which more than half are from sweden (30 in halmstad and 13 in falkenberg)) marked with diamonds scaled to the number of individual reports, and from the swedish national seismic network (snsn) (16 reports) – marked with dots in sweden – are included. the insets in both a and b are all known, instrumentally recorded events over magnitude 3.5 in the two areas. a 2010-02-19 intensity 2 3 4 5 6 57˚n 8˚e t magnitude 5.0 4.5 4.0 3.5 56°n 10°e kattegat fyn sjælland sweden jylland 2012-08-06 b 2 13 7 2 30 3 3 4 2 6 intensity 2 3 4 5 6 12°e57˚n 56˚n magnitude 5.0 4.5 4.0 3.5 56°n 10°e anholt 4444 earthquakes in the area over magnitude 3.5 (see also inset in fig. 4 a). earlier earthquakes are also known historically (gregersen et al. 1991). geus received 344 macroseismic reports with observations of the 2010 earthquake from the public. the earthquake was mainly felt in north-western denmark, with a few reports from northern sjælland and fyn. all observations were classified according to the european macroseismic scale (grünthal 1993), and ranged from 2 to 6, including three instances of slight damage to houses in the form of cracks in walls. kattegat, 6 august 2012, magnitude 4.1 early morning at 02:57 utc (03:57 local time) on 6 august 2012, a magnitude 4.1 earthquake occurred in kattegat, 26 km from the island of anholt, at a depth of 22 km. the focal mechanism indicates a dextral strike-slip movement in a nw–se direction, aligning with the general fault direction in the area including the tornquist zone. also in this area earthquakes are known, both historically (gregersen et al. 1991) and instrumentally recorded (events over magnitude 3.5 in table 1 and inset in fig. 4b). geus received 441 macroseismic reports with observations of the 2012 earthquake from the public. the earthquake was felt mainly in northern sjælland – where the population density is high, and where many people also experienced the magnitude 4.8 earthquake in southern sweden in 2008 (voss et al. 2009). but the earthquake was also felt in north-eastern jylland and northern fyn – and of course on anholt. furthermore, this event was widely felt in sweden, and on fig. 4b observations from snsn (snsn 2012) and usgs (usgs 2012) in usa are included. all observations were classified according the european macroseismic scale (grünthal 1993), and ranged from 2 to a single occurrence of intensity 6 where small cracks had opened in the façade of a house. acknowledgement reynir bodvarsson at university of uppsala kindly supplied us with the macroseismic reports in sweden from the kattegat event. references geus 2012: seismological service. registrerede jordskælv, http://www. geus.dk/departments/geophysics/seismology/seismo_reg-dk.htm gregersen, s., korhonen, h. & husebye, e.s. 1991: fennoscandian dynamics: present-day earthquake activity. tectonophysics 189, 333–344. gregersen, s., hjelme, j. & hjortenberg, e. 1998: earthquakes in denmark. bulletin of the geological society of denmark 44, 115–127. grünthal, g. (ed.) 1993: european macroseismic scale 1992 (updated msk scale). pp. luxemborg: european seismological commision, subcommision of engineering seismology, working group macroseismic scales. lehmann, i. 1956: danske jordskælv. bulletin of the geological society of denmark 13, 88–103. nnsn 2012: norsk nasjonalt seismik nettverk, http://www.geo.uib.no/ seismo/nnsn/index.shtml snsn 2012: svenska nationella seismiska nätet, http://snsn.geofys.uu.se/ thybo, h., sandrin, a., nielsen, l., lykke-andersen, h. & keller, g.r. 2006: seismic velocity structure of a large mafic intrusion in the crust of central denmark from project estrid. tectonophysics 420, 105– 122. usgs 2012: national earthquake information center – neic, http:// earthquake.usgs.gov/regional/neic/ voss, p.h., larsen, t.b., ottemüller, l. & gregersen, s. 2009: earthquake in southern sweden wakes up denmark on 16 december 2008. geological survey of denmark and greenland bulletin 17, 9–12. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: tdj@geus.dk table 1. all instrumentally recorded events over 3.2 on the richter scale in the areas around the north sea and kattegat earthquakes y/m/d-t position depth stamagnitude (utc) (degree) (km) tions (ml) north sea area 1954 / 10 / 18 16:44 56.85n 8.29e 25.5 4 4.6 1964 / 07 / 14 05:33 57.03n 7.20e 36.0 4.0 1969 / 04 / 05 19:09 57.16n 6.76e 0.1 23 4.2 1975 / 11 / 12 06:00 57.10n 7.14e 40.0 37 3.7 1978 / 04 / 26 12:32 56.75n 7.81e 40.0 18 3.4 1981 / 09 / 06 04:11 57.03n 6.88e 40.0 84 5.2 1981 / 09 / 07 14:03 57.06n 7.12e 30.3 13 3.6 1982 / 05 / 24 03:10 56.64n 8.21e 41.4 25 3.7 1987 / 03 / 01 06:42 57.00n 6.98e 40.1 26 3.5 1997 / 11 / 15 16:11 56.86n 7.62e 6.6 56 3.6 1997 / 12 / 04 22:03 56.91n 7.69e 8.5 48 3.4 2001 / 06 / 02 00:44 56.80n 7.80e 59.3 39 3.5 2010 / 02 / 19 21:08 56.89n 7.52e 22.0 111 4.3 kattegat area 1970 / 03 / 12 16:05 56.54n 12.69e 0.0 12 4.0 1980 / 01 / 21 07:41 56.27n 12.16e 10.2 29 3.9 1982 / 11 / 01 02:48 56.28n 11.82e 3.4 17 3.5 1985 / 06 / 15 00:40 56.61n 12.19e 9.1 44 4.7 1986 / 04 / 01 09:56 56.54n 12.18e 7.2 37 4.1 1990 / 05 / 24 09:51 56.48n 11.93e 10.0 15 3.2 1995 / 10 / 04 20:49 56.78n 12.08e 9.9 31 3.8 1997 / 09 / 20 14:21 56.94n 12.42e 15.0 4 3.8 2012 / 08 / 06 02:57 56.60n 11.95e 22.1 31 4.1 mailto:sg@geus.dk geological survey of denmark and greenland bulletin 28, 2013, 13-16 13 geology, seismic activity and groundwater conditions at six potential disposal sites for radioactive waste from risø, denmark peter gravesen, bertel nilsson, merete binderup, tine b. larsen and stig a. schack pedersen in 2003 it was decided by the danish parliament that low and intermediate-level radioactive waste from the danish nuclear research facility, risø, is to be stored at a permanent disposal site (ministeriet for sundhed og forebyggelse 2008; gravesen et al. 2012a). both the issue of storage and the selection of potential sites have caused considerable public debate. in this paper we report on the most recent geological investigations intended to further improve the data base for the selection and decision process, although no conclusions have been drawn at this stage. the waste might be deposited on the land surface, partly below surface or totally below surface, and pre-quaternary rocks and deposits are the main general targets. in 2011 six potential areas were selected for further studyd and evaluation of their suitability for disposal of radioactive waste. the study evaluated local groundwater conditions and earthquake hazards, infrastructure, wildlife, environment, water supply (especially drinking water) and heritage monuments (gravesen et al. 2011, 2012b). the new studies of the six potential sites and their surroundings were performed by the geological survey of denmark and greenland aided by the danish nature agency (naturstyrelsen), and included drilling of several new boreholes. geological setting the areas that were chosen for further study were selected from 22 initial targets (gravesen et al. 2011) based on compliance with specific criteria (see below) that had previously been established from geological models (gravesen et al. 2012a). the six selected areas (fig. 1) are located in different geological settings. the østermarie-paradisbakkerne area on bornholm consists of precambrian paradisbakke migmatite and bornholm gneiss. in the rødbyhavn area on lolland, palaeocene clay rests on maastrichtian chalk, while the kertinge mark area on fyn comprises thick palaeocene clay deposits overlying danian limestone. in the hvidbjerg area of north-west jylland, palaeocene and eocene clay forfig. 1. map of denmark showing the locations of the six areas selected for detailed studies of their suitability as disposal sites for radioactive waste from risø. © 2013 geus. geological survey of denmark and greenland bulletin 28, 13–16. open access: www.geus.dk/publications/bull kerkertinge mark rødbyhavn skive vest hvidbjerg tthise lol land falster møn østermarieparadisbakkerne north sea skagerrak kattegat baltic sea fig. 2 bornholm sjæl land fyn jy l land denmark germany sweden 50 km 55°n 10°e 57°n 14°e 1414 mations rest on danian limestone and are covered by oligocene and miocene clay formations. farther to the east and south-east, at thise and skive vest (west), oligocene and miocene clay, respectively, are present. quaternary clayey till overlies the pre-quaternary rocks and sediments at all sites. collection of seismic, borehole and groundwater data part of the new study consisted of an evaluation of earthquake potential. the seismic activity in denmark has been recorded instrumentally since 1929, and in later years by an interconnected network of seismic stations in denmark and surrounding countries. older earthquakes were examined from descriptions in the literature. for the three potential sites on bornholm, lolland and fyn available seismic data within a radius of 50 km were evaluated. for the three areas in north-west jylland a radius of 75 km was used. in the new study additional boreholes were drilled in each of the six selected areas. sediment samples were collected from each metre and analysed, and several types of continuous geophysical logs were performed in most boreholes. such logs are particularly useful for extrapolations between the samples. on bornholm, eight existing water abstraction wells for domestic use were logged for idenfication of fractures in the bedrock. evaluation of groundwater, drinking water and water supplies was based on new information and existing literature, as well as local knowledge of potential groundwater problems. the latter was obtained by interviewing local specialists and acquisition of written accounts and other information collected by the local authorities. survey results data from the borehole samples and from the geophysical logs have provided important new information on the geological models of the investigated areas, which comprise data on lithology and stratigraphy as well as tectonic and sedimentological structures. all the relevant information has been compiled into six technical reports by the geological survey of denmark and greenland, which are available at geus’ website, http://www.geus.dk/program-areas/natureenvironment/denmark/radioaktivt_affald/index-dk.htm. the main topics are: boreholes and logging, seismic activity and earthquakes, groundwater and drinking water, climate and climate changes and local planning (infrastucture, wildlife, environment, cultural heritage, raw materials, agriculture and water supply structure). all this information will be used in the ongoing evaluation of the areas and identification of the two most suitable areas. in this paper only selected results are presented, including geological traverses through the six areas. geological data the precambrian crystalline rocks on bornholm are thick and extensive and contain a network of fractures. knowledge of the fracture systems in the østermarie-paradisbakkerne area is important for evaluation of groundwater flow and storage, and for evaluation of the stability of the host rocks. horizontal and vertical fractures to a depth of 12 m can be observed in quarries. the number of horizontal fractures appears to decrease downwards, while their spacing increases. this is a normal observation in the uppermost crust. the geophysical borehole logs comprise natural gamma-ray, resistivity, sound velocity, fluid temperature, fluid conductivity and flow logs, which can be used to document fractures at various levels from 20 m below the surface down to a depth fig. 2. geological cross section through palaeogene–neogene deposits of the skive vest area, extending from rettrup in the west to skive in the east. the positions of the boreholes with dgu numbers are indicated along the top of the figure. west (rettrup) d ep th (m ) vejle fjord formation skive claybrejning formation 1 km +40 –30 0 –20 –10 +10 +20 +30 east (skive) 55.1068 .1072 .383 .311 .418 .657 .1228 .1069 .1022 .1229 .533 .503 .532 .152 .85 .180.114 .148a .1120 .507 .1119 palaeogene–neogene deposits clayey till meltwater sand http://www.geus.dk/program-areas/nature-environment/denmark/radioaktivt_affald/index-dk.htm http://www.geus.dk/program-areas/nature-environment/denmark/radioaktivt_affald/index-dk.htm 15 of 90 m. extrapolations of the borehole data suggest that some fractures are horizontal or subhorizontal and connected over larger distances. quaternary clayey and sandy tills on bornholm are mostly only a few metres thick. the pre-quaternary rocks and deposits throughout the remaining parts of denmark are fine-grained, have low permeability, and are covered by quaternary clayey tills. the clay deposits are generally relatively thick and appear to be of wide horizontal distribution within the selected areas, reflecting the fact that the deposits represent marine depositional environments. furthermore, the deposits are characterised by low degrees of glaciotectonic disturbance and pre-quaternary tectonic movements. these characteristics meet the criteria described by gravesen et al. (2011, 2012a) based on the guidelines of the international atomic energy agency (iaea 1994). at rødbyhavn, palaeocene fine-grained deposits and plastic clay from the æbelø and holmehus formations resting on maastrichtian chalk have been documented from geophysical surveys and boreholes drilled for bentonite exploration and geotechnical purposes related to the femern bælt fixed link. the characteristics of the c. 80 m thick clay sequence and the overlying up to 40 m thick hard quaternary clayey till demonstrate that the deposits are in situ, although tectonic disturbances are known to exist outside the area. within and just outside the kertinge mark area, up to 75  m thick palaeocene kerteminde marl and æbelø formation overlying danian limestone have been documented. the new boreholes demonstrate that non-calcareous black clay of the æbelø formation occurs above calcareous grey clay of the kerteminde marl in the northern part of the kertinge mark peninsula, forming the natural stratigraphical sequence. quaternary clayey till up to 40 m thick is found above these formations. some layers are very sandy and gravelly as shown by samples from the boreholes and from geophysical logs. in the hvidbjerg area, palaeocene and eocene grey clay of the kerteminde marl and diatomite from the fur formation rest on danian limestone on the northern flank of the uglev salt diapir. the formations are covered by oligocene and miocene black-brown mica clay deposits from the brejning and vejle fjord formations, with a total of up to 100 m thickness. thick quaternary clayey till comprises the top 20 m. the new borehole demonstrates a more complex buildup of the succession with unconformities caused by tectonic episodes and erosion. in the thise area, more than 100 m thick layers of oligocene mica clay from the brejning formation and green-grey, fine-grained clay from the branden clay unit are covered by up to 18 m thick quaternary clayey till. the new borehole penetrated 22 m of brejning formation below 19 m of quaternary clayey till. the oligocene deposits form an erosional inlier bordered by buried valleys filled with quaternary till and meltwater deposits. in the skive vest area, up to 100 m of oligocene and miocene deposits of fine-grained grey skive clay unit (to the east) and black-brown mica clay from the brejning and vejle fjord formations (to the west) are present; the layers dip slightly to the west (fig. 2). two new boreholes penetrated the two youngest formations down to 50 m below the surfig. 3. earthquake epicentres in the north-west jylland region. the magnitude on the richter scale is shown by the size of the dots (from gravesen et al. 2012b). other earthquake epicentres in the danish region are not shown on this map. magnitude 1 magnitude 2 magnitude 3 magnitude 4 no magnitude aalborg aarhus horsens d e n m a r k kristiansand 58˚n 57˚n 57˚n 56˚n 11˚e9˚e8˚e7˚e6˚e 9˚e 10˚e 11˚e north sea skagerrak kattegat 10˚e 1616 face, possibly indicating that the two formations are found in situ as also indicated by other borehole data in the area. a thin cover of quaternary clayey till is present. seismic activity and earthquakes the seismic activity in most of denmark is low, but a little higher in north-west jylland than in other parts of the country. the most seismically active region relevant for the current investigation is located in skagerrak and the north sea, while epicentres on land are rare (fig. 3). the earthquake magnitudes are low, with values mainly below 3 on the richter scale. no epicentres are found on bornholm, although some occur in the nearshore areas of the baltic sea. bornholm has not been seismically active in recent historical time, although an earthquake in 1875 possibly had its epicentre on west bornholm. small earthquakes in the baltic sea and southern sweden can sometimes be registered on bornholm. according to instrumental and historical sources, lolland does not seem to be seismically active, and the seismicity on fyn is very low. some minor earthquakes have been measured on fyn, but the historical literature does not suggest any earthquake activity. a slightly higher seismicity is found in north-west jylland around hvidbjerg, thise and skive vest, since many small earthquakes located in the north sea and skagerrak are registered on land. historical literature also suggests many small earthquakes around these three areas, and some damage to buildings has occurred. an earthquake in 1841 is probably the largest earthquake registered in denmark. groundwater conditions and drinking water all six areas are situated outside areas of special drinking water interests (osd), which are the most important danish supply areas for drinking water. major regional groundwater reservoirs are also absent, although some small reservoirs are important for local supplies. on bornholm the basement rocks form a fractured groundwater reservoir with restricted groundwater flow at various levels. the reservoir has a limited yield but is used locally in an area where connection to public water utilities is difficult and costly. the rødbyhavn area is almost totally devoid of groundwater reservoirs, although some surface water is extracted for drinking purposes. the kertinge mark area also lacks groundwater reservoirs, but local aquifers close to the surface are used for water supply. just south-east of this area, groundwater reservoirs in an osd area supply the town of kerteminde with drinking water. the reservoir is small and vulnerable because of its restricted size and thickness, and no alternative sources exist in the neighbourhood. in the hvidbjerg area, almost no goundwater is pumped because groundwater reservoirs are lacking. the thise area has some local abstraction from thin meltwater sand reservoirs. an area which is particularly vulnerable to nitrate contamination (nfi) occurs east of thise waterworks. just east and north of the area, osd areas with sand reservoirs are found in buried valleys. throughout the skive vest area, local water supplies use sand and gravel reservoirs; no public or private water utilities are found. final remarks the present investigations of the geology, earthquake risks, groundwater and drinking water conditions are crucial for the selection of two areas for further examination before a final disposal site can be chosen. before the selection of the two areas an environmental impact assessment (eia) has to be performed. also the possibilities of export of the waste to a foreign country have to be investigated. finally the design and establishment of a temporary disposal for the waste have to be described and evaluated. acknowledgement financial support was provided by the parliament of denmark. references gravesen, p., nilsson, b., pedersen, s.a.s. & binderup, m. 2011: lowand intermediate level radioactive waste from risø, denmark. location studies for potential disposal areas. report no. 11. områdebeskrivelser – description of areas. geological survey of denmark and greenland report 2011/51, 64 pp. gravesen, p., binderup, m., nilsson, b. & pedersen, s.a.s. 2012a: geological characterisation of potential disposal areas for radioactive waste from risø, denmark. geological survey of denmark and greenland bulletin 23, 21–24. gravesen, p., nilsson, b., binderup, m. larsen, t. & pedersen, s.a.s. 2012b: lavog mellem radioaktivt affald fra risø, danmark. omegnsstudier. rapport nr. 5. område thise, skive kommune. geological survey of denmark and greenland report 2012/127, 81 pp. iaea 1994: siting of near surface disposal facilities. safety guides. safety series 111-g-3.1, 37 pp. ministeriet for sundhed og forebyggelse 2008: beslutningsgrundlag for et dansk slutdepot for lavog mellemaktivt affald, 47 pp. københavn: ministeriet for sundhed og forebyggelse. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: pg@geus.dk mailto:es@geus.dk geological survey of denmark and greenland bulletin 28, 2013, 21-24 21 late glacial to early holocene development of southern kattegat carina bendixen, jørn bo jensen, ole bennike and lars ole boldreel the kattegat region is located in the wrench zone between the fennoscandian shield and the danish basin that has repeatedly been tectonically active. the latest ice advances during the quaternary in the southern part of kattegat were from the north-east, east and south-east (larsen et al. 2009). the last deglaciation took place at c. 18 to 17 ka bp (lagerlund & houmark-nielsen 1993; houmark-nielsen et al. 2012) and was followed by inundation of the sea that formed a palaeo-kattegat (conradsen 1995) with a sea level that was relatively high because of glacio-isostatic depression. around 17 ka bp, the ice margin retreated to the øresund region and meltwater from the retreating ice drained into kattegat. over the next millennia, the region was characterised by regression because the isostatic rebound of the crust surpassed the ongoing eustatic sea-level rise, and a regional lowstand followed at the late glacial to holocene transition (mörner 1969; thiede 1987; lagerlund & houmark-nielsen 1993; jensen et al. 2002a, b). major parts of kattegat are characterised by thick successions of late weichselian and holocene sediments (mörner 1969; bergsten & nordberg 1992; gyldenholm et al. 1993). at around 9.6 ka bp, a large lagoon–estuary environment in southern kattegat was partly blocked by transgressive, coastal barrier islands and spits (bennike et al. 2000; jensen et al. 2002a). the aim of this paper is to describe the late glacial and early holocene development of southern kattegat, based on a recent study (bendixen 2012). the study area covers 1696.5 km2 and is located south of the island of anholt in the southern part of kattegat (figs 1, 2). the south-western part of the area is shallow but water depths increase to the north-east where depths over 40 m are found (fig. 2). two distinct submarine channels running nearly n–s and ne– sw were probably formed by subglacial meltwater erosion; these channels are partly filled by late glacial and holocene sediments. the area can be seen as a transitional shallow water area at the entrance to the baltic sea (bennike et al. 2000). detailed 2d seismic work will be conducted in the region over the next years, which will improve the basis for interpretations in the coming years. methods the data used in this study consist of shallow single-channel seismic profiles and sediment cores. the seismic data comprise boomer data acquired by r/v alexander von humboldt from 1997 to 1999 and sparker data acquired in 2011 using m/v laura. navigation was based on differential gps. the sediment cores were collected with a 6 m long vibrocorer fig. 1. map of denmark showing the location of the study area of fig. 5 in southern kattegat. fig. 2. bathymetric map of the study area in southern kattegat showing the two incised palaeo-channels from storebælt (arrows) that drained into the kattegat and the location of the shallow seismic profile shown in fig. 3. the white areas to the east are swedish territorial waters. © 2013 geus. geological survey of denmark and greenland bulletin 28, 21–24. open access: www.geus.dk/publications/bull 10°e 14°e 55°n 57°n kattegat storebælt baltic sea kielerkieler fig. 2 fig. 5 øresund fennoscandian shield danish basin 50 km jylland fyn djursland anholt 11°e 12°e 56°40´n 56°20´n anholt djursland 20 km fig. 3 fig. 5 0 10 20 30 40 50 60 70 80 water depth (m) 2222 from laura in 2011. the cores were cut into 1 m long sections that were shipped to the geological survey of denmark and greenland where they were split, photographed, described and subsampled in the laboratory. prior to interpreting the seismic profiles, promax seismic data processing software was used to optimise the data quality. the boomer data were subjected to frequency filtering and the sparker data were subjected to the kirchhoff time-migration method. interpretation of the seismic data was carried out using the program seisvision. results and discussion two late glacial units (lg1 and lg2) have been identified (fig. 3); they show high variability in thickness in the area studied. the sediments were deposited during a sea-level highstand period, which can be seen from the internal seismic pattern. the older lg1 unit shows draping parallel reflections of low amplitude whereas the lg2 unit shows parallel reflections of high amplitude (fig. 3). the units are divided by an erosional unconformity in the north-eastern parts of the study area, whereas continued deposition occurred in the west where no unconformity is found. a single radiocarbon dating of a shell of hiatella arctica from lg1 gave an age of 16.1–16.6 cal. ka bp, and dating of shells from lg2 gave ages of 13.3–15.5 cal. ka bp (jensen et al. 2002a). distinct normal faults cut the late glacial deposits. detailed interpretations of the seismic profiles show that faulting occurred during the last stage of the deposition of the lg2 unit. this is evident because the uppermost part was not affected by the faulting whereas the lower-lying sediments are cut by the faults. this finding is consistent with the conclusions of jensen et al. (2002b). the nw–se-orientated sparker profile r3_021a (fig. 3) shows major bounding faults that cut the late glacial sediments to the north-west and south-east and hence limit the distribution of the younger sediments. within the late glacial units, two major faults are interpreted as strike-slip faults. the faulting postdates the uppermost part of the lg2 unit and was possibly a result of the deglaciation of the katfig. 3. a: selected part of shallow seismic profile r3_021a obtained by a boomer. b. preliminary interpretation. for location see fig. 2. compaction fault normal fault older late glacial sediments lg1 younger late glacial sediments lg2 late glacial lowstand sediments early holocene sediments h1 holocene sediments h2 glacial sediments nw se 30 40 50 60 depth (m) sea floor multiple 30 40 50 60 500 m sea floor multiple a b deformed deformed 23 tegat region that led to isostatic uplift and reactivation of older faults. this tectonic event may have contributed to the opening of storebælt. the late glacial units between the faults show an internal pattern with contorted reflectors. the late glacial deposits have been reworked by faulting and a significant erosional unconformity is found between the late glacial and the holocene sediments. this unconformity formed during the late glacial – early holocene lowstand period. above the erosion surface, two holocene units can be separated on the basis of their difference in reflection pattern and infill direction, with h1 showing infill from the south-east and h2 from the north-east. two lithological units presumably of holocene age were also found in sediment core no. dgu 561118.10 collected at 56°23.165́ n, 11°22.0´e from a water depth of 38.0 m (fig. 4). the core was 518 cm long and consisted of 136 cm sand with abundant shells of the common blue mussel mytilus edulis that is characteristic of shallow water, overlain by 328 cm of mud with shells of turritella communis, arctica islandica, pecten s.l. sp. and other marine molluscs that are characteristic of deeper water. the marked lithological change 382 cm below the core top probably corresponds to the boundary between units h1 and h2 (bendixen 2012). we suggest that the sand was deposited during the early holocene when sea level was low, whereas the mud was deposited after the relative sea level had increased. radiocarbon dating of holocene sub-littoral sand deposits in the region has yielded ages of c. 11–10 cal. ka bp (bennike et al. 2000; jensen et al. 2002a). a palaeogeographic map of the region illustrates northward coastal progradation with spits and barriers with backbarrier-enclosed environments in which finer-grained sediments were deposited (fig. 5). fig. 4. sedimentological log of vibrocore dgu 561118.10. for location see fig. 5. fig. 5. palaeogeographic map of the area south of anholt in kattegat in early holocene showing glacial deposits to the south and west, coastal sandy deposits formed by prograding shorelines and a barrier island. the present-day form of anholt is included to show the location of the map. mud cla y sil t vf sand f m c vcli th ol og y thin clay layer with organic material carbonate present carbonate present arctica islandica cerastoderma, turritella dark grey colour a few thin clay layers pecten s.l. sp. dark olive grey colour olive grey colour turretella communis abundant mytilus edulis a few stones up to 3 cm 100 200 300 400 500 d ep th b el ow c or e to p (c m ) glacial deposits coastal deposits sea profile r3_021a (fig. 3) barrier island core 561118.10 10 km anholt 11°49´e 56°31´n 56°40´n 2424 conclusions the late glacial sediments in southern kattegat consist of a lower and an upper sequence deposited during relatively high sea level; the boundary between the sequences shows an erosional surface towards the north-east. the distribution of the sediments is limited by major faults which were initiated during deposition of the uppermost part of the youngest late glacial unit. major faults bounding the late glacial sediments were active during the deposition of the uppermost part of the youngest late glacial unit. we suggest that strike-slip movements occurred due to isostatic reactivation of the fennoscandian border zone and upward movement of the late glacial sediments. an early holocene lowstand level is identified as an erosional surface, underlying units h1 and h2. initial transgression resulted in coastal progradation and back-barrier-enclosed environments with deposition of finergrained sediments (h2) in the former incised valleys. acknowledgement the danish nature agency (naturstyrelsen) funded the work. references bendixen, c. 2012: interpretation of shallow seismic and sediment cores from the area stretching from the southern part of kattegat to the great belt in the period late-weichselian to early holocene 1, 2, 74 pp + 102 pp. unpublished m.sc. thesis, university of copenhagen, denmark. bennike, o., jensen, j.b., konradi, p.b., lemke, w. & heinemeier, j. 2000: early holocene drowned lagoon deposits from the kattegat, southern scandinavia. boreas 29, 272–286. bergsten, h. & nordberg, k. 1992: late weichselian marine stratigraphy of the southern kattegat, scandinavia: evidence for drainage of the baltic ice lake between 12,700 and 10, 300 years bp. boreas 21, 223–252. conradsen, k. 1995: late younger dryas to holocene palaeoenvironments of the southern kattegat, scandinavia. the holocene 5, 447– 456. gyldenholm, k.g., lykke-andersen, h. & lind, g. 1993: seismic stratigraphy of the quaternary and its substratum in southeastern kattegat, scandinavia. boreas 22, 319–327. houmark-nielsen, m., linge, h., fabel, d., schnabel, c., xu, s., wilcken, k.m. & binnie, s. 2012: cosmogenic surface exposure dating the last deglaciation in denmark: discrepancies with independent age constraints suggest delayed periglacial landform stabilisation. quaternary geochronology 13, 1–17. jensen, j.b., kuijpers, a., bennike, o. & lemke, w. 2002a: balkat – the baltic sea without frontiers (english version). geologi, nyt fra geus 2002(4), 20 pp. jensen, j.b., petersen, k.s., konradi, p., kuijpers, a., bennike, o., lemke, w. & endler, r. 2002b: neotectonics, sea-level changes and biological evolution in the fennoscandian border zone of the southern kattegat sea. boreas 31, 133–150. lagerlund, e. & houmark-nielsen, m. 1993: timing and pattern of the last deglaciation in the kattegat region, southwest scandinavia. boreas 22, 337–347. larsen, n.k., knudsen, k.l., krohn, c.f., kronborg, c., murray, a.s. & nielsen, o.b. 2009: late quaternary ice sheet, lake and sea history of southwest scandinavia: a synthesis. boreas 38, 732–761. mörner, n.-a. 1969: the late quaternary history of the kattegatt sea and the swedish west coast. déglaciation, shorelevel displacement chronology, isostasy and eustasy. sveriges geologiska undersökning ser. c, 63(3), 487 pp. thiede, j. 1987: the late quaternary skagerrak and its depositional environment. boreas 16, 425–432. authors’ addresses c.b., j.b.j., o.b. & l.o.b.*, geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: cb1@geus.dk * also: department of geosciences and natural resource management, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. mailto:jbj@geus.dk palaeomagnetic results from the lopra-1/1a re-entry well, faroe islands 51© geus, 2006. geological survey of denmark and greenland bulletin 9, 51–65. available at: www.geus.dk/publications/bull palaeomagnetic results from the lopra-1/1a re-entry well, faroe islands niels abrahamsen the palaeomagnetic dating and evolution of the faroe islands are discussed in the context of new density and rock magnetic results from the deepened lopra-1/1a well. the reversal chronology of the c. 6½ km thick basalt succession is also described. the polarity record of the faroe islands may now be correlated in detail with the geomagnetic polarity time scale. the lowermost (hidden) part of the lower basalt formation correlates with chron c26r (selandian age), the top (exposed) part of the lower basalt formation correlates with chrons c26n, c25r and c25n (selandian and thanetian age) and the middle and upper basalt formations correlate with chron c24n.3r (ypresian). inclinations indicate a far-sided position of the palaeomagnetic poles, which is characteristic of results from most palaeogene volcanics from the northern north atlantic region. the density, magnetic susceptibility and magnetic remanence of 20 specimens from one solid core (1½ m in length) and 26 sidewall cores from the well between –2219 and –3531 m below sea level (b.s.l.) suggest that the volcanic materials can be divided into two characteristic groups: solid unaltered basalts and altered basalts and tuffs. the magnetic properties are typically log-normally distributed and the carriers of remanence are ti-poor ti-magnetites with curie temperatures close to 580°c. the inclination of the 1½ m core at 2380 m b.s.l. is dominantly negative (two plugs at the very top of the core do show normal polarity, but they are likely to be misoriented as all specimens appear to be from one flow). magnetic logging (magnetic susceptibility and field intensity) down to 3515 m b.s.l. was made in lopra-1/1a together with other geophysical logs but did not yield conclusive inclination data. keywords: palaeomagnetism, rock magnetism, magnetic reversals, plate tectonics, faroe islands, lopra-1/1a well, north atlantic, large igneous province _________________________________________________________________________________________________ department of earth sciences, university of aarhus, finlandsgade 8, dk-8200 aarhus n, denmark. e-mail: abraham@geo.au.dk review of the geology the faroe islands are situated on the eastern side of the northern north atlantic between the shetland islands and iceland on the northern part of the ne–sw-trending elongated faroe rise (fig. 1). the volcanic islands are a result of the hotspot-related plume activity recorded by the britoarctic large igneous province (lip) (lawver & muller 1994; larsen & saunders 1998; t.b. larsen et al. 1999; burke & torsvik 2004) that stretches from present-day central west greenland to the north-western parts of the uk. seismic (e.g. richardson et al. 1998) and gravity investigations (e.g. saxov & abrahamsen 1964) suggest that the invisible basement of the islands is composed of continental lithospheric crust, somewhat thinned by lithospheric stretching processes during the continental breakup that formed the north atlantic. the exposed part of the faroe islands is composed of a c. 3 km thick pile of palaeogene flood basalts (rasmussen & noe-nygaard 1969, 1970; noe-nygaard & rasmussen 1984) situated above a c. 3½ km unexposed volcanic sequence below sea level (fig. 2). only minor sedimentageus bulletin no 9 7 juli.pmd 07-07-2006, 14:1951 52 ry layers are intercalated in the whole volcanic sequence. the volcanic sequence, more than 6½ km in total thickness (waagstein 1988; l.m. larsen et al. 1999), is divided into three parts, the lower (> 4½ km thick), the middle (1.4 km thick) and the upper basalt (> 0.9 km thick) formations. the basalts are cut by numerous dykes and a few sills. an up to 10 m thick coal-bearing formation of lacustrine claystones and shales was deposited on top of the slightly eroded surface of the lower basalt formation (the a-level). two other stratigraphical levels, b (in the middle formation) and c (separating the middle and upper formations), are also useful for stratigraphical purposes (fig. 2). the purpose of the present contribution is to present 60° norway 50° 50° 30° 30° 15° 15° 0° 0° 15° 15° 30° 30° 30°w 30°w 15°w 15°w 0°w 0°w icelandd icelandd rockall rockall plateau plateau 70° 70° 50° 60° 70 50° 70° 30° 15° 0° 15° 30° c al ed on ia n fr on t 30°w greenland norway faroe islands 15°w c al ed on ia n fr on t iceland uk ridg e dk 0°w onshore basalt and sills offshore basalt flows and sills seaward-dipping reflector sequences rockall plateau spr ea din g r idg e fig. 1. index map of the eastern north atlantic showing the faroe islands (modified from larsen et al. 1995). n2.2 c25n c24n.3n c24n.1n c23n.2n c28n c29n magnetic polarity chrons (1) 52 54 56 58 50 48 ma 60 62 64 66 68 70 c26n c25n c22n c21n c24n.3n c24n.1n c23n.2n c27n c28n c29n c30n c31n c32n.1n faroe islands stratigraphy (4) –4.5 km lower fm middle fm upper fm 0 sea level 1 2 3 4 5 6 2.9 km d an ia n yp re sia n s el an di an t ha ne tia n r1 r1 r2 r2 r3 n3 n3.2 n2.1 n3.1 n3r n2 n2.2 n2r c24r c25r c26r (3) reversals (2) fig. 2. compilation of magnetic reversals within the c. 6½ km thick basalt pile of the faroe islands, showing stratigraphy and the correlation with the geomagnetic polarity time scale. the four columns are based upon information compiled from: (1) ogg (1995); (2) abrahamsen (1965, 1967), abrahamsen et al. (1984), waagstein (1988) and riisager et al. 2002a; (3) tarling & gale (1968); (4) rasmussen & noe-nygaard (1970), waagstein (1988) and l.m. larsen et al. (1999). geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1952 53 the magnetic results from a core of basaltic rock obtained from the lopra-1/1a reentry well and to discuss these data in relation to other palaeomagnetic results from the faroe islands. previous work magnetic investigations in relation to the faroe islands have been made since the early 1960s (abrahamsen 1965, 1967; saxov & abrahamsen 1966; tarling & gale 1968; tarling 1970; schrøder 1971; løvlie 1975; løvlie & kvingedal 1975; abrahamsen et al.1984; schönharting & abrahamsen 1984; tarling et al. 1988). density determinations (saxov & abrahamsen 1964) as well as gravity measurements (saxov 1969) and seismic investigations (pálmason 1965; bott et al. 1974, 1976; casten 1974; nielsen et al. 1981; richardson et al. 1998) have been made on and around the islands. geophysical logs from the lopra-1/1a and vestmanna-1 boreholes have been published by nielsen et al. (1984), boldreel (2006, this volume) and abrahamsen & waagstein (2006, this volume). geothermal measurements were described by balling et al. (1984) and balling et al. (2006, this volume). palaeomagnetic results from the faroe islands have been published by abrahamsen (1965, 1967), tarling & gale (1968), tarling (1970), løvlie (1975), løvlie & kvingedal (1975), abrahamsen et al. (1984), schönharting & abrahamsen (1984) and riisager et al. (2002a, b). a comparison of palaeomagnetic results from east greenland and other results from the palaeogene of the north atlantic igneous province (naip) was published by tarling et al. (1988) and a critical review of palaeomagnetic poles from the eurasian part of the naip together with a new pole for the faroe islands was presented by riisager et al. (2002a). a summary of all published palaeomagnetic directional data from the faroe islands is shown in table 1. the magnetic results for the exposed part of the basalt succession were extended by the wells at vestmanna-1 and lopra-1 in 1980–1981 (abrahamsen et al. 1984; schönharting & abrahamsen 1984) and by the re-entry of the lopra-1/1a hole in 1996, the results of which are presented in this paper. despite intentions, the re-entry hole at lopra-1/1a reached a depth of 3565 m without penetrating to the base of the lower basalt formation volcanics. the polarity sequence and the compiled total stratigraphic column of the faroe islands as now known are shown in fig. 2, together with the geomagnetic polarity time scale (gpts). essentially we find three intervals of reverse magnetic polarity (r1, r2, r3) with two normal table 1. palaeomagnetic results from the faroe islands all formations all formations ubf, torshavn ubf, argir mbf, argisfossar mbf, vestmanna core lbf, vestmanna core mbf + lbf, vestm. core mbf + lbf lbf, lopra-1; 862 mbf lbf, lopra-1; 1219 mbf lbf, lopra-1; 1923 mbf lbf, lopra-1; 2178 mbf lbf, lopra-1a; 2380 mbf average, nos 1–10, except * average, nos 1, 2, 3 & 6 n(dg) 33 1809 34 8 18 275 28 303 548(43) 6 8 5 7 20 (10) (4) 176.0 185.0 171.9 175.1 156.0 7.7 181.2 ± 69.0 ± 66.4 –72.2 –53.9* –36.0* –61.8 ± 63.4 ± 61.9* ± 60.9 –75.0 –62.0 –73.0 –55.0* –71.7 ± 67.2 ± 67.3 6 1.9 3.5 2.2 1.2 6.3 1.2 4.5 1 3 2.0 1.4 6.3 53 258 53.4 19.4 46.1 24.5 709 213 80.0 76.7 84.0 62.3 48 70.8 72.8 70.9 71.4 78.7 plat °n 159.0 161.0 218.0 182.4 154.7 164.4 plon °e 10.3 3.1 6.3 3.1 1.9 10.0 1.9 6.0 8.0 a 95 ° 52.5 48.9 57.3 34.4* 20.0* 43.0 45.0 43.1* 41.9 61.8 43.2 58.6 35.5* 56.5 50.0 50.1 palaeolat (°n) 1 2 3 * * 4 5 * 6 7 8 9 * 10 11 12 no. r&n r&n r r r r r&n r&n r&n r ?r r ?r r r&n r&n (1) (2) (3) (3) (4) (5) (5) (5) (6) (7) (7) (7) (7) (8) (8) (8) referenceformation/site/core depth decl ° incl ° k 14 δ95 ° α 95 ° polarity lbf, mbf, ubf: lower, middle and upper basalt formations; no.: number in palaeolatitude figure; * not used in average; n: number of samples; dg: directional groups; decl: mean of cleaned declination; incl: mean of cleaned inclination; α95 : cone of 95% confidence. for core data the inclination statistics of kono (1980) were used for k and α 95 (tarling 1983); k: fisher precision parameter; plat: latitude of apparent palaeomagnetic pole; plon: longitude of apparent palaeomagnetic pole; δ95 and a 95 : error angles of app. latitude and app. palaeopole at 95% confidence level; (1) abrahamsen 1967; (2) tarling 1970; (3) løvlie & kvingedahl 1975; (4) løvlie 1975; (5) abrahamsen et al. 1984; (6) riisager et al. 2002a; (7) schönharting & abrahamsen 1984; (8) this work. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1953 54 polarity intervals in between (n2 and n3). minor differences between columns (2) and (3) in fig. 2 are likely to be due to somewhat different positions of the profiles investigated on suðuroy, the southernmost of the faroe islands. according to recent high-precision 40ar-39ar datings (storey et al. 1996; l.m. larsen et al. 1999), the basalt formations in the faroe islands as well as the contemporaneous east greenland basalts can be divided into an older part with ages of about 59–56 ma, followed (after a pause or a period with much reduced volcanic activity) by a younger part, with ages of 56–55.5 ma for the faroes and 56– 54.5 for east greenland. based upon these radiometric datings, the polarity record of the faroe islands may now be correlated to the gpts as shown in fig. 2. the lowermost (hidden) part of the lower basalt formation correlates with chron c26r (selandian age), the upper (exposed) part of the lower basalt formation correlates with chrons c26n, c25r and c25n (selandian and thanetian) and the middle and upper basalt formations correlate with chron c24r (ypresian). this correlation follows the suggestion by waagstein (1988), who revised the original interpretation of abrahamsen et al. (1984) by suggesting that r3 belongs to chron c26r rather than to c24r. more details in relation to magnetic inclinations from the lopra-1/1a data are discussed below. assuming the geomagnetic field to have been a central, axial dipole field, the palaeolatitude may be determined from the characteristic (primary) inclination of the volcanics, combining both polarities. a compilation of all inclination values obtained from the faroe islands is listed in table 1. using inclination statistics (kono 1980; tarling 1983) the fisherian mean of published inclinations (group numbers 1–10, table 1) is 67.2° ± 1.4° (equivalent to a palaeolatitude of 50.0° ± 2.1°), whereas the average of the palaeolatitudes listed is 50.9° ± 2.3° (± 1 sigma). further discussion of the shallow inclinations and the palaeolatitude question will be given below. palaeogeography many palaeogeographic reconstructions of the north atlantic have been published since the early work of bullard et al. (1965) (e.g. ziegler 1990; knott et al. 1993; l.m. larsen et al. 1999; torsvik et al. 2001; mosar et al. 2002). before about 60 ma, the supposed mantle hotspot (just south-east of iceland at the present day) lay under the volcanic areas of disko and nuussuaq in west greenland (o’connor et al. 2000; nielsen et al. 2002; chambers et al. 2005), far from the faroe islands that are situated just north-west of the continental margin of europe. the whole volcanic pile of the faroe islands, more than 6½ km thick, was formed in the time interval between chron 26 (61.65 ma) and chron 24n.3n (c. 53.286 ma) (chron ages are the orbitally tuned age calibration of gradstein et al. 2004, table 5.2). during this time interval, the hotspot moved eastwards under greenland as greenland moved westnorth-west relative to europe and the north atlantic gradually opened between the faroe islands and greenland. absolute declinations are known from only four of the palaeomagnetic investigations from the faroe islands (ta1 3 6 21 3 6 n 21 3 6 fig. 3. palaeomagnetic directions from the faroe islands (from table 1, results nos 1, 2, 3 and 6) with α95 circle. the axial dipole field direction is indicated by a cross. fig. 4. apparent palaeomagnetic pole positions (solid circles) with 95% significance circle (table 1, poles nos 1, 2, 3 and 6). all poles appear ‘farsided’ as seen from the faroe islands (diamond). further discussion in the text. 6 2 1 3 6 6 1 1 3 3 6 2 1 3 geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1954 55 ble 1, nos 1, 2, 3 and 6). if both normal and reverse polarities are combined and assumed to be normal directions towards the north with steep down-dip (positive) inclinations, the four directions appear as shown in fig. 3. the equivalent apparent palaeomagnetic pole positions are shown in fig. 4. all four poles are seen to be ‘farsided’ (wilson 1971; merrill et al. 1998), the apparent palaeomagnetic poles falling beyond the geographic pole as seen from the faroe islands. an equivalent histogram of all published apparent palaeolatitudes (table 1) is also illustrated in fig. 5, most of which show low values as compared to the present-day latitude. lopra-1/1a investigations sampling, instruments and techniques used the material investigated from the extended lopra-1/1a well consists of two types: core plugs from the solid core (2380.0 to 2381.3 m) and sidewall cores (between 2219 and 3531 m). the solid core is 1.4 m long and in several pieces, but some fit together, as shown in fig. 6. after marking the core with an upward directed arrow in the core lab at geus, 20 plugs with a diameter of 2.5 cm were drilled orthogonal to the main core and cut to a standard length of 2.2 cm. the major part of the present magnetic investigation is concentrated upon these 20 plugs. in addition some rotary sidewall cores were investigated. the pieces from the sidewall cores had a diameter of 2.33 cm and varied in length, which limited the possibility of fitting these samples into the magnetic instrument holders. bulk density to avoid problems with air bubbles adhering to the relatively small specimens if they were weighted in water, the bulk density was determined by weighing in air only (to an accuracy of ± 0.001 g), then determining the volume by measuring the shape of the specimens (to an accuracy of ± 0.02 cm). the likely accuracy in the finally determined density is about ± 2–3%, depending on the rough0 20 40 60 80 result no. (table 1) pa la eo la tit ud e 1 2 3 4 5 6 7 8 9 10 present latitude fig. 5. histogram of apparent palaeolatitudes according to table 1 (poles nos 1–10). most results appear systematically low compared to the present-day latitude. for further discussion see the text. 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 5 10 15 20 25 30 35 40 45 50 55 60 65 70 75 80 85 90 95 100 105 110 115 120 125 fragments 130 cm fragments fragments fragments 0 fig. 6. sketch from a photograph of the lopra-1/1a core between 2380 and 2381.4 m. the fragments containing the numbered 1inch core plugs are indicated. the absolute azimuths of the individual segments and fragments are not known. the top segment containing plugs nos 1 and 2 appears to have been turned upsidedown before the core was archived. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1955 56 ness of the shape. a total of 34 specimens were determined (table 2). susceptibility two types of susceptibility instruments were used. initial whole core measurements were made by a handheld czech kappameter kt5 (sensitivity ± 0.00001 si) before drilling plugs from the core. a molspin bulk susceptibility bridge (sensitivity ± 0.000001 si) was then used to measure the susceptibility of the core plugs and to monitor possible chemical changes during thermal demagnetisation experiments. a total of 46 specimens were measured (table 2). remanence the direction (declination and inclination) and intensity of the natural remanent magnetisation (nrm) was determined using a molspin spinner magnetometer. the plugs with preferred dimensions of 2.2 cm in length and a diameter of 2.5 cm (plugs from the solid core) were all measured and demagnetised in detail, see below. the nrm of the sidewall cores was also measured but, due to the variable length of the core pieces, only one (swc57) was investigated in detail (table 2). the sensitivity of the molspin spinner is ± 0.02 ma/m and the direction of the remanence within the plug was determined to within ± 1°. the declination and inclination is given with respect to the local specimen coordinates, assuming the axis of the plug (= specimen) to be approximately horizontal (i.e. orthogonal to the lopra-1/1a drill hole). as the azimuth of the vertical core is not known, the true magnetic decli1v 1h 1 2 3 4 5 6 6v 6h 7 8 9 10 11 12 13 14 15 16 17 18 19 20 20v 20h 2380.08 2380.08 2380.08 2380.12 2380.17 2380.21 2380.28 2380.58 2380.58 2380.58 2380.62 2380.66 2380.69 2380.75 2380.79 2380.91 2380.94 2380.97 2381.00 2381.05 2381.09 2381.12 2381.22 2381.25 2381.25 2381.25 basalt basalt basalt basalt basalt basalt basalt basalt basalt basalt basalt basalt basalt basalt basalt basalt basalt basalt basalt basalt basalt basalt basalt basalt basalt basalt 69.040 36.000 31.79 28.66 30.38 28.78 29.76 28.22 38.240 44.907 31.37 26.24 28.84 29.92 27.31 29.37 30.07 27.08 26.96 30.99 29.11 30.24 29.87 25.87 71.848 27.740 2.992 2.981 2.885 2.868 2.922 2.869 2.92 0.05 0.02 6 17997 3258 327 1247 673 2319 2553 1227 1391 971 1539 333 1782 1628 984 359 887 322 2616 2705 2256 3808 851 20 table 2a. lopra-1/1a: magnetic susceptibility, nrm, q-ratio, density solid core (d = 25 mm) sample no. depth m rock type weight g nrm corr ma/m density g/cm3 46.26 41.62 17.32 3.903 6.945 4.340 11.773 18.850 11.459 16.865 33.756 34.636 32.987 37.825 36.450 38.704 43.323 31.722 35.241 41.288 36.890 43.15 28.40 14.00 2.98 22 susceptbridge 10–3 si 10.868 4.728 2.105 4.513 3.897 4.951 3.404 2.691 2.073 0.723 1.117 0.254 1.184 1.123 0.639 0.208 0.703 0.230 1.592 1.843 2.44 2.51 0.56 20 q-ratio nrm/(f × sus) a: mean standard deviation mean error n geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1956 57 nation is not known. a further description of the palaeomagnetic experimental standard laboratory procedures may be found in e.g. butler (1992). af-demagnetisation after measuring the initial nrm intensity, 16 of the 20 plugs were af-demagnetised in stepwise increasing alternating magnetic fields (table 3) using a molspin af-demagnetiser. minimum and maximum af-fields were 2.5 mt (25 oe) and 100 mt (1000 oe), respectively. thermal demagnetisation stepwise thermal demagnetisation was made in a schonstedt furnace on four plugs from the solid core. an initially moderate af-demagnetisation of up to between 7.5 and 15 mt removed recently induced viscous magnetisation components, most likely acquired during the drilling operations (details in table 2), after which the thermal demagnetisation was applied. b: a & b: 1.43 0.55 0.71 0.64 0.61 0.69 1.60 1.83 0.72 1.13 0.62 0.64 0.83 0.39 0.58 0.58 0.88 77.4 0.623 4.864 0.032 0.046 0.060 0.015 1.406 1.240 9.224 5.022 2.344 2.030 1.080 2.346 0.074 0.470 0.024 2.433 *) excluding sample no. 59 59 57 46 44 43 40 39 38 37 36 34 33 31 30 26 25 19 16 15 13 12 9 6 5 5 4 2219.00 2275.00 2441.00 2456.00 2475.00 2558.00 2559.80 2560.20 2562.00 2570.00 2610.00 2630.00 2690.00 2780.00 2970.00 3030.00 3233.50 3328.00 3382.00 3438.00 3464.50 3500.50 3512.50 3514.50 3514.50 3531.00 sample no. depth m basalt basalt basalt basalt, ves. basalt, ves. basalt, ves. basalt, ves. tuff, lapilli tuff basalt basalt, alt. tuff, lapilli tuf, lapilli basalt tuf, lapilli basalt tuff basalt basalt tuf, lapilli tuf, lapilli basalt tuff tuf, lapilli tuf, lapilli basalt rock type 18.966 22.222 23.857 33.653 18.421 17.973 22.920 21.428 29.838 9.353 18.741 13.720 8.082 21.170 18.744 18.515 9.915 18.471 14.066 19.505 17.287 16.494 23.953 13.372 15.648 27.341 weight g 2.361 2.857 2.865 2.488 1.981 2.708 2.711 2.336 2.345 2.379 3.167 2.404 2.276 2.929 2.553 2.808 2.572 2.875 2.919 2.549 2.605 3.001 2.648 2.511 2.607 2.940 density g/cm3 4406 2180 9027 0.7 1.3 1.5 0.4 38.5 3412 587 365 7171 304 1146 106 1.8 12.0 0.8 38.2 nrm ma/m 1.43 87.96 46.65 0.55 0.71 0.64 0.61 0.69 69.17 1.60 1.83 76.89 3.77 26.67 0.72 1.13 63.36 0.62 0.64 0.83 0.39 0.58 0.58 0.88 susceptbridge 10–3 si 87.96 46.65 69.17 76.89 3.77 26.67 63.36 high > 2 × 10–3 10–3 si low < 2 × 10–3 10–3 si q-ratio nrm/(f × sus) mean standard deviation mean error n mean standard deviation mean error n 2.63 0.27 0.05 26 2.68 0.27 0.05 32 1516 2585 593 19 1895 3276 525 39 16.20 28.80 5.88 24 34.47 21.00 3.90 29 34.5 21.4 4.0 29 0.85 0.39 0.10 17 1.85 2.34 0.55 18*) 2.16 2.45 0.40 38*) table 2b. lopra-1/1a: magnetic susceptibility, nrm, q-ratio, density sidewall cores (d = 23.3 mm) 22.1 23.6 3.5 46 geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1957 58 the limit between the two groups being about 2 × 10–3 si. the higher group yields an average susceptibility of 34 ± 4 (σ = 21) × 10–3 si and the lower group an average susceptibility of 0.85 ± 0.1 (σ = 0.39) × 10–3 si. the more strongly magnetised group is represented mostly by unaltered basalts, whereas the less strongly magnetised group is more typical of most sediments including tuffaceous sediments, as well as vitrinites and deuterically altered or weathered basalts. in the present case the difference between high and low values in the basalts is likely to be caused by alterations of the primary ti-magnetites, since ti-magnetite is the main carrier of the remanence (see below). nrm intensity the intensity of the nrm (natural remanent magnetisation) is listed in table 2 and shown in fig. 7. values of the rev.: norm: all rev: mean inclination standard deviation n mean inclination standard deviation n mean inclination standard deviation n –71.22 1.99 18 76.00 5 2 –71.70 2.85 20 α 95 = 1.95 k = 709 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 2380.08 2380.12 2380.17 2380.21 2380.28 2380.58 2380.62 2380.66 2380.69 2380.75 2380.79 2380.91 2380.94 2380.97 2381.00 2381.05 2381.09 2381.12 2381.22 2381.25 (?n) (?n) r r r r r r r r r r r r r r r r r r polarity 31.79 28.66 30.38 28.78 29.76 28.22 31.37 26.24 28.84 29.92 27.21 29.37 30.07 27.08 26.96 30.99 29.11 30.24 29.87 25.87 weight g 0–40 0–40 0–70 0–7.5 0–95 0–40 0–70 0–40 0–50 0–630 0–50 0–600 0–50 0–40 0–50 0–630 0–70 0–630 0–50 0–50 treatment af mt/°c 18.00 3.26 0.33 1.25 0.67 2.32 2.55 1.23 1.39 0.97 1.54 0.33 1.78 1.63 0.98 0.36 0.89 0.32 2.62 2.71 335 350 358 175 345 51 62 75 80 289 106 7 169 204 195 334 163 334 32 24 81 71 –70 –70 –69 –67 –71 –69 –71 –75 –72 –72 –70 –71 –70 –74 –73 –74 –73 –71 6 2 2 4 1 1 1 1 1 3 2 2 2 1 2 2 3 2 3 4 nrm intensity a/m decl (rel.) characteristic direction mad degrees demag interval af thermal mt tmax °c 15–50 5–40 15–70 10–95 5–40 5–70 5–40 5–50 5–50 5–50 5–40 5–50 10–70 5–50 0–50 t630 t600 t630 t630 sample no. depth m incl table 3. lopra-1/1a: palaeomagnetic results results bulk density the bulk density of sidewall cores and plugs are listed in table 2. the mean bulk density of the solid basaltic core was well-determined as 2.92 ± 0.02 (standard deviation (σ) = 0.05) g/cm3, although the determination was based on only six specimens. the bulk density of the swc-cores was lower and much more scattered, 2.63 ± 0.05 (σ = 0.27) g/cm3. the low bulk density and high scatter is likely to be due to differences in porosity and the abundance of secondary minerals. susceptibility the magnetic susceptibility is also listed in table 2 and shown in fig. 7. in contrast to for instance the bulk density, the magnetic susceptibility and remanence intensity may vary considerably, and they are known typically to be logarithmically normal distributed (e.g. tarling 1983; abrahamsen & nordgerd 1994). this is also the case here, as two log-normal distributions are found. the core susceptibility varies between 4 and 46 × 10–3 si, and the susceptibility of the swc-cores varies even more, between 0.4 and 88 × 10–3 si. the data thus fall into two populations, geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1958 59 order of 1 a/m are typical for the unaltered basalts, whereas tuffs and altered basalts may have lower values. the mean nrm intensity value of the core plugs is 2.26 ± 0.85 (σ = 3.8) a/m, and the mean nrm intensity value of the swccores is 1.52 ± 0.59 (σ = 2.6) a/m. due to the high scatter, the mean nrm intensities of the two groups are not significantly different. the combined populations plotted logarithmically (fig. 9) again show two overlapping lognormal distributions, as do the susceptibilities (fig. 8). q-ratio the q-ratio (koenigsberger ratio) illustrated in fig. 10 is the ratio between the remanent (jnrm) and the induced (ji = k·f) magnetisation, q = jnrm/ji = jnrm/(k·f), f being the intensity of the local geomagnetic field, f ≈ 0.05 mt. for basaltic rocks, values between 0.2 and 10 are characteristic. generally, the higher value the more fresh and unaltered the samples are. mean values for q is found to be 2.4 ± 0.6 (σ = 2.5) for the core plugs, and 1.8 ± 0.6 (σ = 2.3) for the swc-cores, respectively (omitting a single extraordinary high value of q = 77 for swc59). the q-ratios of the two groups are not significantly different, but again the combined population has a tendency to two log-normal distributions. magnetic carriers two examples of isothermal remanent magnetisation (irm) acquisition of plugs nos 3 and 7 are shown in fig. 11. both specimens show magnetic saturation around 0.1 t, which indicate that the dominant carrier of the remanence is magnetite or ti-magnetite, although maghemite may also be present. the thermal demagnetisations (see below) show blocking temperatures between 560 and lopra-1: susceptibility frequency (si) 0 2 4 6 8 10 log (susceptibility × 10–3 si) n = 46 lopra-1: nrm frequency 0 2 4 6 8 10 log (nrm, ma/m) n = 39 lopra-1: q-ratio frequency 0 2 4 6 8 10 log (q-ratio) n = 38 –0.5 0.0 0.5 1.0 1.5 2.0 –1 0 1 2 3 4 5 –2 –1 0 1 2 n um be r n um be r n um be r fig. 7. histograms of susceptibility, nrm intensity and q-ratio. all appear bimodal on a logarithmic scale. 2200 2400 2600 2800 3000 3200 3400 3600 lopra-1: swc and core 0 1 10 100 depth (m) core su sc ep tib ili ty ( 10 –3 s l) fig. 8. magnetic susceptibility of sidewall cores (diamonds) and core plugs (dotted line shows extent), logarithmic scale. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1959 60 580°c, indicating that the ti-content is low, pure magnetite having a curie temperature of 580°c (e.g. dunlop & özdemir 1997). af and thermal demagnetisations the nrm values of the 20 cores and the swc-cores are listed in tables 2a & b, and examples of characteristic results of the af and thermal demagnetisation experiments performed are illustrated in fig. 12. thermal demagnetisations were made on cores nos 10, 12, 16 and 18 and af-demagnetisations were made on the remaining 16 cores. values chosen for the af-field were in most cases 0, 5, 7.5, 10, 15, 20, 25, 30, 40, 50 and 60 mt. some plugs were further demagnetised to 70, 80, 90 and 100 mt. four thermally demagnetised plugs were first af-demagnetised in 2.5, 5, 7.5, and 10 mt fields, to remove the recent drillstem-induced viscous remanence (see below), and then stepwise demagnetised at temperatures of 150, 250, 350, 450, 550, 570, 600 and 630°c. the examples in fig. 12 show stereographic plots (left) of the direction of the unit vector of the remanent magnetisation (solid signature: positive inclination, open signature: negative inclination). all except the first example show characteristic stable negative inclinations. to the right, the corresponding intensity decay of the sample is shown (normalised to the initial value j0 = jnrm), the horizontal scale indicating the peak value of the applied alternating field in oe (× 0.1 mt), or the temperature in c. inclination prior to the demagnetisation experiments, about half of the plugs showed a low coercivity nrm with positive inclination (down-dip), which is most likely due to a drillstem induced viscous remanent magnetisation (vrm). the vrm was easily removed by af-demagnetisation in low fields, typically between 2.5 and 5 mt. based upon the afand thermal demagnetisation data, the characteristic (stable) remanent magnetisations for each plug were determined by the principal component analysis (pca) method of kirschvink (1980), as implemented in the iapd-programme by torsvik (1986). in all cases a stable characteristic, supposed primary, mag0 1 10 100 1000 100 000 10 000 core n r m ( m a /m ) depth (m) 2200 2400 2600 2800 3000 3200 3400 3600 lopra-1: swc and core fig. 9. nrm intensity of side-wall cores (diamonds) and core plugs (dotted line shows extent), logarithmic scale. lopra-1: swc and core 0 2 4 6 8 10 12 14 16 18 20 depth (m) core 77.4 q -r at io * 2200 2400 2600 2800 3000 3200 3400 3600 * q-ratio = (nrm/f × susceptibility) fig. 10. q-ratio of side-wall cores (diamonds) and core plugs (dotted line shows extent). geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1960 61 netisation was isolated, as listed in table 3, and illustrated in the stereogram of fig. 13. (bearing in mind that the azimuth of the core is not known, only the inclinations are diagnostic, declinations being relative.) after cleaning, plugs nos 3 to 20 show typical steep negative inclinations. only plugs 1 and 2 show normal inclinations and they are both from the topmost 10 cm long core-piece. the broken core is from the massive centre of a very thick flow and it is most unlikely that the inclination should shift the sign within the core. it is therefore suggested that the top part of the core has been turned upside-down, most likely during the initial handling at the core site. as the azimuth of the core is not known, ordinary fisher statistics are not applicable, but the modified inclination statistics of kono (1980) may be used. supposing all 20 plugs to have negative inclinations, the mean value is found to be: im = –71.7°, with α95 = 1.95° (k = 709, n = 20) provided that the drilling was truly vertical. this would give an unusually accurate determination of the palaeofield inclination. however, as the geomagnetic secular variation cannot be recorded from one flow only, this low value irm lopra-1 0 100 200 300 400 500 600 700 0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1 field (t) lopra-1 plug 7 lopra-1 plug 3 m ag ne tic in te ns ity ( a /m ) 1 0 0 100 mt j/jo j/jo 1 0 0 100 mt 1 0 0 100 mt 1 0 0 100 mt 1 0 0 100 mt 1 0 0 100 mt af af af + th af + th af af j/jo j/jo j/jo j/jo sample: l1-2 sample: l1-14 sample: l1-3 sample: l1-18 sample: l1-10 sample: l1-20 fig. 11. isothermal remanent magnetisation (irm) of plugs nos 3 and 7. both specimens show magnetic saturation around 0.1 t, indicating that the dominating carrier of the remanent magnetisation is magnetite. fig. 12. examples of typical behaviour of samples during af and thermal demagnetisation. plugs nos 2 (af 0–40 mt), 3 (af 0–70 mt) and 10 (af 0–10 mt, combined with heating up to 630°c). plugs nos 14 (af 0– 40 mt), 18 (af 0–10 mt, combined with heating up to 630°c) and 20 (af 0–50 mt). geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1961 62 of α95 does not give a realistic estimate for the accuracy of the average palaeomagnetic field inclination. furthermore an angle of c. 2.5° from the vertical towards the southeast has been obtained from the ghmt-log of the hole at 2380 m (r. waagstein, personal communication 2005). the inclination i0 of the geocentral axial dipole (gad) field at the site of lopra/suðuroy with a latitude of 61.4°n is equivalent to a value of i0 = 74.8°, which is about 3° steeper than that found for the core. based on mcelhinny & mcfadden’s (1997) analysis of a large number of volcanic data from the last 5 ma in the global palaeomagntic database, the expected geomagnetic dispersion of a vgp (virtual geomagnetic pole) at the latitude of the faroe islands may further be estimated to be c. 20°. earlier palaeomagnetic investigations have typically given systematically lower mean values (see figs 3, 5) for the palaeomagnetic inclination (table 1) except the one of –72° for a site near torshavn (løvlie & kvingedal 1975). if we suppose the value from the lopra-1/1a core of im = –71.7° to be the optimum one, this would correspond to an axial dipole palaeolatitude for the lopra-1/1a site of 56.2°n at the time of extrusion. most cenozoic palaeopoles tend to be ‘farsided’ (wilson 1971; merrill et al. 1998), i.e. the palaeofield recorded in the rocks shows a more shallow inclination than does the present-day geomagnetic field at the site and biased shallow inclinations are also the case for most of the palaeogene volcanic palaeomagnetic data from the north atlantic region. this phenomenon may be due either to northward plate tectonic movements after the formation of the sample, non-symmetric behaviour of the geomagnetic field at the time of formation or unusual magnetic properties of the rocks investigated – or a combination of all three effects. a systematic error due to the latter cause (magnetic refraction) is not likely, as this requires rather strong values of the magnetic properties of the lavas (e.g. knudsen et al. 2003). if the palaeogeomagnetic field was exactly a geocentral axial dipole field (the gad-hypothesis), this would imply that the lithospheric plate carrying the faroe islands had moved about 5.2° northward during the last c. 60 ma with an average northward component of velocity of c. 1 cm/year. an octopole contribution of the order of 10% (i.e. g3°/g1° = 0.1) to the central axial dipole field would alone suffice to explain the observed farsidedness of the faroe islands. an octopole contribution of this order of magnitude has been considered for precambrian and palaeozoic as well as mesozoic times (e.g. kent & smethurst 1998; torsvik et al. 2001; van der voo & torsvik 2001). however, merrill & mcfadden’s (2003) analysis of data in the palaeomagnetic global database for the last 5 ma concluded that a nondipole bias appears less likely for the younger periods. therefore, rather than claming that the shallower value of inclination indicates fully either a northward plate movement of 5.2° (the ‘traditional’ palaeomagnetic interpretation), or is due entirely to a deficiency in the gadhypothesis, a more cautious interpretation may be a combination of both, implying that the gad-hypothesis may not be exactly valid for the early palaeogene, i.e. that the palaeomagnetic field was not a perfect central and axial dipole field at that time. to solve this palaeomagnetic important question fully, more global data from the period is needed. reversal stratigraphy and age at lopra-1/1a as mentioned above, the reversal stratigraphy of the 6½ km thick faroe basalt formations was re-interpreted by waagstein (1988), based upon published data then available (abrahamsen 1967; tarling & gale 1968; schönharting & abrahamsen 1984; abrahamsen et al. 1984), including the former palaeomagnetic results from lopran ee fig. 13. characteristic af-cleaned inclinations of the 20 plugs from the lopra-1/1a core at depths between 2380 and 2381.4 m. plugs other than nos 1 and 2 (with positive inclinations, solid symbols) have negative inclinations (open symbols). the declinations are arbitrary since the azimuth of the core is not known. full circle shows the expected axially centred dipole inclination of 74.7° at the lopra-1/1a drill site, i.e. 3° steeper than the numerical average of –71.7° (dashed circle) of the 20 core plugs (see table 1). geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1962 63 1, the cored information of which at that time reached a depth of 2178 m. all five cores from lopra-1, at depths of 338, 862, 1219, 1923 and 2178 m, showed negative inclinations, i.e. reversed polarity, although no stable values were obtained after demagnetising the cores from 338 and 1219 m (schönharting & abrahamsen 1984). the bottom of lopra-1 was interpreted by waagstein (1988) to match marine anomaly 26r, thus superseding two earlier alternative correlations discussed by abrahamsen et al. (1984), in which this level was suggested to match either marine anomaly 25r or 24r. the present data from lopra-1/1a, with negative inclinations in the single core from 2380 m depth, indicates a reversed polarity at this level. provided that there are no reversals in the unsampled interval above, the present data extend the reversed sequence of the lower basalt formation from the core at td of the original well (2178 m) to the present level of the solid core at 2380 m. the swccores reach the deeper level of 3531 m. however, as the up–down orientation of the individual swc-cores is not known, no inclination information has yet been obtained from below 2381 m. combining all polarity evidence available from the faroe islands and comparing with the paleocene time scale by berggren et al. (2000), we conclude that the lower part (below sea level) of the lower basalt formation may be correlated with chron c26r (selandian age), while the upper (exposed) part of the lower basalt formation correlates with chrons c26n, c25r and c25n (selandian and thanetian age). the middle and upper basalt formations correlate with chron c24r (ypresian age). magnetic logging (magnetic susceptibility and field intensity) was also attempted in the lopra-1/1a well together with other geophysical logs (boldreel 2006, this volume) but, due to technical problems with the magnetic logging tool, no reliable inclination data were obtained (abrahamsen & waagstein 2006, this volume). summary and conclusions a compilation of the palaeomagnetic age, the reversal chronology and evolution of the c. 6½ km thick basalt formations of the faroe islands is presented, together with new petrophysical results from the lopra-1/1a well. 1. the polarity record of the faroe islands has been correlated in detail with the global polarity time scale. the lower part (below sea level) of the lower basalt formation correlates with chron c26r (selandian age).the upper (exposed) part of the lower basalt formation correlates with chrons c26n, c25r and c25n (selandian and thanetian age). the middle and upper basalt formations correlate with chron c24r (ypresian age). 2. the inclinations yield farsided positions for the palaeomagnetic poles, which is characteristic of most palaeogene volcanics and sediments from the north atlantic region. 3. the density and the rock magnetic properties of a solid core (1½ m in length) and 26 sidewall cores from the lopra-1/1a well between –2219 and –3531 m are bimodal and suggest two characteristic groups of volcanic materials, solid unaltered basalts and altered basalts and tuffs. 4. the magnetic properties are typically log-normally distributed and the carriers of remanence appear to be ti-poor ti-magnetites with curie temperatures close to 580°c. 5. the inclination of the 1½ m core at –2380 m is predominantly negative. 6. magnetic logging of magnetic susceptibility and field intensity was made in lopra-1/1a down to –3515 m together with other geophysical logging, but yielded inconclusive inclinations. acknowledgements the rock and palaeomagnetic measurements were made in the geophysical laboratory of the department of earth sciences, university of aarhus. informative discussions with regin waagstein and the access to material from the lopra-1/1a well at geus are acknowledged. suggestions for improvements of the manuscript from regin waagstein, john piper, the editor and an anonymous referee are also acknowledged. references abrahamsen, n. 1965: geofysiske undersøgelser på færøerne, 144 pp. unpublished cand. scient. thesis, aarhus universitet, danmark. abrahamsen, n. 1967: some palaeomagnetic investigations in the faeroe islands. bulletin of the geological society of denmark 17, 371–384. abrahamsen, n. & nordgerd, p. 1994: rock magnetism of tertiary volcanics from north-east greenland. rapport grønlands geologiske undersøgelse 162, 195–200. abrahamsen, n. & waagstein, r. 2006: magnetic logs from the lopra-1/1a and vestmanna-1 wells, faroe islands. geological geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1963 64 survey of denmark and greenland bulletin 9, 41–49 (this volume). abrahamsen, n., schönharting, g. & heinesen, m. 1984: palaeomagnetism of the vestmanna-1 core and magnetic age and evolution of the faeroe islands. in: berthelsen, o., noe-nygaard, a. & rasmussen, j. 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(eds): early tertiary volcanism and the opening of the ne atlantic. geological society special publication (london) 39, 225–238. wilson, r.l. 1971: dipole offset – the time average palaeomagnetic field over the past 25 million years. geophysical journal of the royal astronomical society 22, 491–504. ziegler, p.a. 1990: geological atlas of western and central europe. amsterdam: elsevier. manuscipt received december 1999; revision accepted 16 june 2005. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1965 geological survey of denmark and greenland bulletin 1, 115-144 115 this paper attempts a further integration and standardisation of the jurassic dinoflagellate cyst zonation schemes established for the british and danish areas (fig. 1; davey 1979, 1982; woollam & riding 1983; nøhr-hansen 1986; riding & thomas 1988, 1992; poulsen 1991, 1992, 1994a; koppelhus & nielsen 1994). in addition, the relationships between the zonation and palaeoecology are discussed, particularly with respect to the appearance and disappearance of species in relation to changes in sea level and palaeotemperature. dinoflagellates dinoflagellates are primarily motile single-celled algae. about half are autotrophic, phytosynthetic species, the reminder being non-photosynthetic consumers that ingest other organisms or particulate organic matter as predators, symbionts, parasites or decomposers. they have been a major component of marine phytoplankton since their diversification in the late triassic and became important in non-marine environments during the early cretaceous (batten & lister 1988). the dinoflagellates have diversified into a wide range of ecological habitats and many of them are sensitive to environmental physical/chemical changes. the understanding of the response of extant dinoflagellates to environmental stress is not well-advanced. the prediction of environment, temperature, water depth and other parameters based on certain species, however, can be made with some accuracy (wall et al. 1977; de vernal et al. 1992). the jurassic dinoflagellate cyst zonation of subboreal northwest europe niels e. poulsen and james b. riding with an appendix by bjørn buchardt: oxygen isotope palaeotemperatures from the jurassic in northwest europe the jurassic dinoflagellate cyst zonation for the british–danish area is revised and discussed in relation to palaeoenvironmental factors, in particular, eustatic changes and fluctuations in palaeotemperature. the stepwise evolution of dinoflagellate cyst assemblages as defined by inceptions and apparent extinctions was largely controlled by sea-level change, particularly during intervals with significant short-term eustatic fluctuations. during times characterised by less pronounced, or longer term, sea-level change, fluctuations in oceanic palaeotemperatures appear to have influenced dinoflagellate evolution. differences in the ranges of certain taxa between denmark and the united kingdom may be partly related to differences in palaeotemperature. keywords: subboreal northwest europe, jurassic, dinoflagellate cyst zonation, palaeotemperatures and biotic provincialism, dinoflagellate palaeoecology n.e.p., geological survey of denmark and greenland, geocenter copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: nep@geus.dk j.b.r., british geological survey, keyworth, nottingham ng12 5gg, uk. b.b., geological institute, university of copenhagen, geocenter copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. geological survey of denmark and greenland bulletin 1, 115–144 (2003) © geus, 2003 dinoflagellate cysts represent the non-motile dormant stage in the life cycle, resulting from sexual fusion (evitt 1985); it is not known, however, whether encystment is related solely to sexual reproduction. cysts are generally resistant to adverse conditions, whereas the motile thecate stage is quickly destroyed after death or encystment. most resting cysts act as sedimentary particles and eventually sink to the sea floor, although many of them have adaptations to floating such as processes, an oil-rich cellular content or becoming entangled with floating debris (sarjeant et al. 1987). the cyst distribution pattern is therefore not only dependent on ecological factors but also on sedimentary history. knowledge of the major factors controlling the distribution patterns of jurassic dinoflagellate cysts cannot be obtained by analogy to modern taxa. however, detailed studies of pre-neogene fossil dinoflagellates during the last 60 years have revealed distinctive distribution patterns of provincialism related to palaeosalinity, palaeotemperature and palaeowater-depth gradients. geological framework during the latest triassic and the earliest jurassic, denmark, the united kingdom, and the north sea basin were part of the subboreal province (fig. 2). the biogeographical affiliations of parts of this area changed, however, as the tethyan realm periodically expanded northwards, and occasionally retreated southwards. the main reasons for this are thought to be variations in seafloor spreading, the migration routes of faunas and floras or a combination of these factors. in general, the true boreal realm consisted of east greenland (jameson land and adjacent areas), the greenland sea (northernmost atlantic), the boreal sea, northern siberia, the arctic islands and northern canada (enay 1972, 1980). the norwegian–danish basin, the british isles and the central graben are termed the subboreal province (fig. 2). this province may be further divided into: (1) a southern part including southern england and the southern central graben, termed the northwest european subprovince, and (2) a northern part, or transitional area, i.e. the 116 poland dknorth sea germanynetherlands atlantic ocean ireland norway sweden england scotland celtic sea wealddorset east midlands london– brabant massif moray firth hebrides skagerrak– kattegat platform yorkshire (cleveland) brora bornholm skånejylland central graben baltic shield ? ?jurassic strata, outcrop/subsurface 500 km norwegian–danish basin 0° 16°e 32°e8°w16°w 58°n 50°n 8°e fig. 1. generalised distribution of jurassic strata in northwest europe, both at outcrop and in the subsurface. compiled from brooks & chesher (1975), czaplicka (1976), michelsen (1978), cope et al. (1980a, b), guy-ohlson (1986), guy-ohlson & norling (1988), ziegler (1988), andrews et al. (1990), cameron et al. (1992), hamblin et al. (1992), rattey & hayward (1993), stoker et al. (1993), erlström et al. (1994), hamann (1994), koppelhus & nielsen (1994), vejbæk & britze (1994) and japsen et al. (2003, this volume). dk, denmark. norwegian–danish basin, northern united kingdom and the northern central graben, referred to as the boreal– subboreal subprovince (fig. 2; enay 1972, 1980). following deposition of the mainly continental triassic succession, jurassic sedimentation was largely marine and clastic-dominated. a number of major depocentres developed during the jurassic in the subboreal area, including the celtic sea, weald, east midlands, cleveland, moray firth, hebrides, central graben and danish– norwegian basins (fig. 1). large parts of the north sea region were affected by the late toarcian – aalenian central north sea thermal or volcanic doming (sellwood & hallam 1974; whiteman et al. 1975; eynon 1981; ziegler 1988; underhill & partington 1993). following resumed subsidence and initial (middle jurassic) paralic/deltaic sedimentation, the north sea region accumulated a thick marine succession dominated by the mud-rich deposits of the kimmeridge clay formation and its correlatives. during the jurassic period, representing some 70 ma of earth history, the sedimentary successions attained (post-compaction) thicknesses in excess of 1600 m in onshore uk (hallam 1992a), over 1200 m in the danish basin (the danish part of the norwegian–danish basin, see fig. 1; nielsen 2003, this volume) and over 4000 m in the deepest parts of the danish central graben (møller 1986; sundsbø & megson 1993; japsen et al. 2003, this volume). methods the ranges for the stratigraphical index dinoflagellate cyst species are based on those recorded in both the british 117 boreal realm tethys 2 month polar winter night volgian toarcian pliensbachian arctic canada g re en la nd se a siberia polish subprovince northwest european subprovince subboreal province boreal–subboreal subprovince present day land land in jurassic times position of jurassic north poles c. 1000 km fig. 2. biogeographical realms, provinces and subprovinces (modified from poulsen 1996, based on enay 1972, 1980). the approximate geographical extent of the pliensbachian and volgian two-month polar night is indicated, assuming that precession was of the same order in the mesozoic as at the present-day. and danish areas. the zonation represents an idealised succession of bioevents, which takes into account factors such as local hiatuses, palaeoenvironmental/facies control on sediment distribution and ranges of taxa, natural variability and sampling problems. range extension using graphical correlation (shaw 1964; edwards 1984, 1989) was not undertaken. significant differences in ranges are not generally observed, although occasional exceptions have been recorded and are indicated on the composite range charts (see figs 3–5). the middle jurassic of the danish central graben and the danish basin consists mainly of coarse-grained, non-marine strata deposited during the aalenian–bathonian regressive event, or is represented by hiatuses. it contains few stratigraphic levels with marine intercalations and dinoflagellate cysts from this part of the middle jurassic are extremely rare (for further discussion, see poulsen 1992, 1996). the dinoflagellate cyst ranges for the middle jurassic are therefore based on ranges determined from the united kingdom. facies control may have influenced the dinoflagellate cyst assemblages, making them variable in composition inter-regionally. the principal aim here is to present a zonation for the subboreal area that is tested against other faunal zonations, such as those for ammonites and ostracods, and in which apparent variations in ranges have been taken into account, thus giving the best overall regional correlations. as a result of variations in sample availability, inadequate data in certain intervals means that the exact correlation of dinoflagellate cyst zonal boundaries to ammonite zonal boundaries may be uncertain. in such cases, the boundaries are deemed to be coincident, although this methodology may introduce certain minor errors. the ammonite zones are chronostratigraphical units (wimbledon & cope 1978; callomon 1984; cox 1990; page 2003, this volume) and are thus referred to by the species name alone, in roman type (e.g. tenuicostatum zone). this is the convention followed by working groups of the international subcommission on jurassic stratigraphy (isjs) and the international commission on stratigraphy (ics). the base of the jurassic is taken at the inception of the ammonite genus psiloceras, which marks the base of the planorbis zone as recommended by warrington et al. (1994; for further discussion, see page 2003, this volume). the dinoflagellate cyst zones are indicated by the generic and specific names of the index taxon in italics, e.g. dapcodinium priscum zone. only the taxa which are useful for identifying each zone or subzone are cited; accessory forms are consistently present but these are not usually biostratigraphically diagnostic. stratigraphic palynology davey & riley (1978) and morbey & dunay (1978) presented a summary of knowledge of upper triassic and jurassic dinoflagellate cyst biostratigraphy in northwest europe. williams & bujak (1985) subsequently published an extensive synthesis of dinoflagellate cyst zonation schemes for the triassic to quaternary interval. riding & thomas (1992) and riding & ioannides (1996) outlined the history of study of jurassic dinoflagellate cyst biostratigraphy. poulsen (1991, 1992, 1993, 1994a, b, 1996, 1998) demonstrated the utility of the british jurassic dinoflagellate cyst zonations of woollam & riding (1983) and riding & thomas (1992) in denmark and poland. this paper attempts further emendation, integration and standardisation of jurassic dinoflagellate cyst zonal schemes. the zonation presented herein (table 1) is an integrated zonation for denmark and the united kingdom based on the zonations presented by davey (1979, 1982), woollam & riding (1983), nøhr-hansen (1986), riding & thomas (1988, 1992) and poulsen (1991, 1992, 1994b, 1996). the british and danish zonations are in general identical, with the exception of certain new subzones and other emendations that resulted from simultaneous, independent revisions by poulsen (1992) and riding & thomas (1992); these differences are integrated and standardised here. only the necessary emendations and other pertinent comments for unifying the british and danish zonations are given below. it should be noted that the middle–upper jurassic dinoflagellate cyst zones have been recognised in poland (poulsen 1992, 1993, 1994a, 1996, 1998). these zones are also widely recorded in the jurassic deposits in europe and adjacent regions, for example in onshore united kingdom (the wessex, east midlands, cleveland, hebrides and onshore moray firth basins), the danish basin and in the danish sector of the north sea central graben. the zones have been named following biostratigraphical tradition and international stratigraphical guidelines and recommendations, with the zonal name related to an index species, for example, the acanthaulax senta zone. following international rules, zonal names change to conform with any valid changes to the name of the index species (hedberg 1976; salvador 1994); for example, the acanthaulax senta zone of woollam & riding (1983) became the liesbergia scarburghensis zone of riding & thomas (1992) and is now the trichodinium scarburghensis zone. further118 119 zonation formal zonal name reference# (this study) dsk2 endoscrinium pharo subzone (or gochteodinia villosa zone, subzone c) 1 (2 6) dsk1 rotospaheropsis thula subzone (or gochteodinia villosa zone, subzone b (pars)) 1 9 (2 6) dsj39 gochteodinia villosa zone, subzone b (pars) 9 (6) dsj38 exmontodinium expiratum subzone (pars) or gochteodinia villosa zone, subzone a 1 6 9 dsj37 dingodinium spinosum zone 1 dsj36 dichadogonyaulax culmula zone, subzone b (pars) 1 2 5 6 9 dsj35 dichadogonyaulax culmula zone, subzone a 1 2 5 6 dsj34 glossodinium dimorphum zone, subzone e 6 dsj33 glossodinium dimorphum zone, subzone d 6 dsj32 glossodinium dimorphum zone, subzone c 6 dsj31 glossodinium dimorphum zone, subzone b 6 dsj30 glossodinium dimorphum zone, subzone a 6 dsj29 endoscrinium luridum zone, perisseiasphaeirdium pannosum subzone 5 dsj28 endoscrinium luridum zone, stephanelytron scarburghense subzone 5 dsj27 scriniodinium crystallinum zone, subzone d 4 6 dsj26 scriniodinium crystallinum zone, subzone c 4 6 dsj25 scriniodinium crystallinum zone, subzone b 4 6 dsj24 scriniodinium crystallinum zone, subzone a (pars) 9 (6) dsj23 scriniodinium crystallinum zone, subzone a (pars) 9 (6) dsj22 trichodinium scarburghense zone, subzone b 9 (2) dsj21 trichodinium scarburghense zone, subzone a 9 (2) dsj20 wanaea fimbriata zone 2 6 dsj19 wanaea thysanota zone 2 6 dsj18 ctenidodinium continuum zone 6 (2) dsj17 ctenidodinium sellwoodii zone, subzone c 6 (2) dsj16 ctenidodinium sellwoodii zone, subzone b 6 (2) dsj15 ctenidodinium sellwoodii zone, subzone a 6 (2) dsj14 cribroperidinium crispum zone, subzone b 6 9 (2) dsj13 cribroperidinium crispum zone, subzone a 6 9 (2) dsj12 nannoceratopsis gracilis zone, subzone e 6 dsj11 nannoceratopsis gracilis zone, subzone d 6 dsj10 parvocysta nasuta zone (stat. nov.) (pars), (nannoceratopsis gracilis zone, subzone c) 9 (5 6) dsj9 parvocysta nasuta zone (stat. nov.) (pars), (nannoceratopsis gracilis zone, subzone b) 9 (5 6) dsj8 mancodinium semitabulatum zone (emend. nov.) 9 (2 5 6) dsj7 luehndea spinosa zone, subzone b 6 dsj6 luehndea spinosa zone, subzone a, or luehndea spinosa zone, subzone b 6 or 8 dsj5 nannoceratopsis senex zone (new), (luehndea spinosa zone, subzone a) 9 (8) dsj4 mendicodinium reticulatum zone, (liasidinium variabile zone, subzone b) 7 (2) dsj3 liasidinium variabile zone, subzone a 2 dsj2* dapcodinium priscum zone, subzone b 2 5 dsj1* dapcodinium priscum zone, subzone a 2 5 dstr rhaetigonyaulax rhaetica zone 2 table 1. jurassic dinoflagellate cyst zonation # references: 1: davey (1978, 1982) 2: woollam & riding (1983) 3: nøhr-hansen (1986) 4: riding & thomas (1988) 5: poulsen (1991, 1992) 6: riding & thomas (1992) 7: koppelhus & nielsen (1994) 8: poulsen (1994a) 9: this study * note that in the united kingdom it is not possible to separate these two zones, hence they are termed zone dsj1–2 more, riding & thomas (1992) used a three-letter abbreviation for their jurassic dinoflagellate cyst zones (e.g. lsc for the liesbergia scarburghensis zone). calcareous nannoplankton, foraminiferal and other biozones are often denoted by an alphanumeric code in which the first letter indicates the respective fossil group, for example n for (calcareous) nannoplankton, and the second letter indicates a period or epoch – p for palaeogene and n for neogene. each zone is thus enumerated np24, np25, nn1, nn2, nn3 etc. this method results in abbreviations (e.g. nn9) which are easier and more convenient to use, especially for nonpalaeontologists; this methodology is adopted here. the dinoflagellate cyst zones evaluated in this study are thus given a similar code e.g. zone dsj13, (d for dinoflagellate cysts, s is added to emphasise that it is a subboreal zonation, tr for triassic, j for jurassic and k for cretaceous, see table 1). the jurassic and cretaceous are numbered, whereas only one zone is recognised for the triassic. as almost every jurassic dinoflagellate cyst zone is divided into subzones, the subzones are generally taken as the basic numbered biounits in this study. the zonation is, where possible, defined both by species with first occurrences and last appearances coincident with the zonal boundaries. dinoflagellate cyst zonation the zonation presented below is an attempt to further refine the work of r.j. davey, h. nøhr-hansen, n.e. poulsen, j.b. riding, j.e. thomas, r. woollam and others, and to integrate and standardise the jurassic dinoflagellate cyst zonation scheme (table 1). accounts of the ranges of the index species can be found in raynaud (1978), davey (1979, 1982), fisher & riley (1980), woollam (1980), riding (1982, 1984a, b, 1987), riley & fenton (1982), woollam & riding (1983), riding et al. (1985), riding & sarjeant (1985), nøhr-hansen (1986), riding & thomas (1988, 1992, 1997), riley et al. (1989), poulsen (1991, 1992, 1993, 1996), riding et al. (1991), partington et al. (1993) and j.b. riding, j.e. thomas and i.p. wilkinson (in: richards et al. 1993). triassic zonation dstr; upper triassic (rhaetian) a single zone, dstr, is established for the upper triassic (rhaetian). the dstr zone corresponds to the rhaetogonyaulax rhaetica zone of woollam & riding (1983), which is based on the common/abundant presence of rhaetogonyaulax rhaetica (fig. 3). this zone may locally be further divided into two or three new zones defined on the abundance, presence and/or absence of sverdrupiella spp., heibergella spp., suessia spp. and rhaetogonyaulax rhaetica. these relatively high diversity associations are most likely to occur in the peripheral regions of the subboreal province. jurassic – lowermost cretaceous zonation dsj1, dsj2; uppermost triassic – lower sinemurian (turneri zone) the dsj1 and dsj2 zones are equivalent to the dapcodinium priscum zone of woollam & riding (1983); the lower part of the dsj1 zone spans the triassic–jurassic boundary. subzones a and b of the dapcodinium priscum zone are herein named the dsj1 and dsj2 zones, respectively (fig. 3). in denmark, the dsj1 zone is coincident with the range of d. priscum in the uppermost triassic and lower sinemurian; the index species is absent in the dsj2 zone. the range top of the index species, d. priscum, appears to become younger in more southerly areas relative to the northern subboreal province. as stated above, in the danish embayment the d. priscum zone may be subdivided into two subzones, the upper of which is characterised by the absence of d. priscum. in the united kingdom, d. priscum has been recorded in the lower sinemurian turneri zone (riding 1984a), thereby negating the bipartite subzonal division of the d. priscum zone in the united kingdom initially advocated by woollam & riding (1983). in the united kingdom, the zones established here are termed the dsj1–2 zones to indicate that the d. priscum zone is not subdivided. further south, in portugal, davies (1985) recorded d. priscum from the lower pliensbachian. dsj3; upper sinemurian (obtusum–raricostatum zones) this zone corresponds to the liasidium variabile zone, subzone a of woollam & riding (1983) and riding & thomas (1992) and coincides with the total range of the index species (fig. 3). 120 dsj4; lower pliensbachian (jamesoni – davoei (pars) zones) the dsj4 zone is characterised by the presence of mendicodinium reticulatum or is marked by the absence of dinoflagellate cysts. koppelhus & nielsen (1994) erected the equivalent lower pliensbachian mendicodinium reticulatum zone below the luehndea spinosa zone. the mendicodinium reticulatum zone is equivalent to subzone b of the liasidium variabile zone of woollam & riding (1983) and riding & thomas (1992). dsj5–7; lower pliensbachian – lower toarcian (davoei (pars) – tenuicostatum zones) the upper pliensbachian luehndea spinosa zone of woollam & riding (1983) was expanded into the lowermost toarcian and subdivided into subzones a and b by riding & thomas (1992) and poulsen (1994a). however, the subzones were defined differently by these authors and subzone b of poulsen (1994a) corresponds to subzone a of riding & thomas (1992). a three-fold division of the luehndea spinosa zone is thus proposed below. dsj5; lower pliensbachian (davoei zone (pars)) the dsj5 zone corresponds to the luehndea spinosa zone, subzone a of poulsen (1994b) and is here formally renamed the nannoceratopsis senex zone. the zone is defined in poulsen (1994a) as the interval from the inception of n. senex to the first occurrences of l. spinosa and other species of nannoceratopsis (fig. 3). dsj6; pliensbachian (margaritatus, spinatum zones) this interval corresponds to the luehndea spinosa zone, subzone a of riding & thomas (1992) and subzone b of poulsen (1994a). the zone is defined as subzone a in riding & thomas (1992) and additional species are used to define the base of subzone b of poulsen (1994b). it is therefore defined herein as the interval from the inceptions of l. spinosa, mancodinium semitabulatum, maturodinium inornatum, nannoceratopsis gracilis, n. raunsgaardii, n. ridingii, n. triceras and valvaeodinium armatum to the range tops of m. inornatum and v. armatum. dsj7; lowermost toarcian (tenuicostatum zone) the dsj7 zone corresponds to the luehndea spinosa zone, subzone b of riding & thomas (1992) and was defined as the interval from the apparent extinction of maturodinium inornatum and valvaeodinium armatum, to the range top of luehndea spinosa (fig. 3). dsj8; lower toarcian (falciferum, bifrons (pars) zones) the dsj8 zone broadly corresponds to the mancodinium semitabulatum subzone of poulsen (1992) and subzone a of the nannoceratopsis gracilis zone of riding & thomas (1992; table 1). the dsj8 zone also forms part of the mancodinium semitabulatum zone, subzone a of woollam & riding (1983). dsj9–10; lower toarcian – lowermost aalenian (bifrons (pars) – opalinum zones) the parvocysta nasuta range subzone of poulsen (1992) is herein raised in status to that of a zone and is formally divided into two subzones, a and b, corresponding to subzones b and c, respectively, of the nannoceratopsis gracilis zone of riding & thomas (1992). the inceptions of nannoceratopsis dictyambonis, phallocysta elongata, and susadinium scrofoides define the boundary between subzones a and b, which are equivalent to the dsj9 and dsj10 zones, respectively (fig. 3). the stratigraphically important species nannoceratopsis ambonis, nannoceratopsis dictyambonis, scriniocassis priscus and scriniocassis weberi are not found in the danish basin, although they are characteristic elements both in the united kingdom (woollam & riding 1983; riding 1987) and germany (prauss 1989). dsj11, dsj12;aalenian – lower bajocian (murchisonae–sauzei zones) the dsj11 and dsj12 zones correspond precisely to subzones d and e, respectively, of the nannoceratopsis gracilis zone of riding & thomas (1992; fig. 4, table 1). dsj13, dsj14; lower–upper bajocian (humphriesianum–parkinsoni zones) the dsj13 and dsj14 zones are equivalent to the acanthaulax crispa zone of riding & thomas (1992). the acanthaulax crispa zone of riding & thomas (1992) 121 122 210 205 200 195 190 185 180 sv er dr up ie lla m ut ab ilis h ei be rg el la k en de lb ac hi a su es sia s w ab ia na va lva eo di ni um d ia cr or ha et iu m be au m on te lla ? ca m in us pi na rh ae to go ny au la x rh ae tic a d ap co di ni um p ris cu m li as id iu m v ar ia bi le m en di co di ni um r et ic ul at um n an no ce ra to ps is se ne x m at ur od in iu m in or na tu m va lva eo di ni um a rm at um lu eh nd ea s pi no sa sc rin io ca ss is w eb er i m an co di ni um s em ita bu la tu m n an no ce ra to ps is gr ac ilis pa rv oc ys ta " su ite " pa rv oc ys ta n as ut a ph al lo cy st a eu m ek es ju ra ss ic t ri as si c la te t ri as si c r ha et ia n h et ta ng ia n si ne m ur ia n ea rl y ju ra ss ic pl ie ns ba ch ia n t oa rc ia n pe ri od ep oc h a ge t im e in m a sc rin io ca ss is pr isc us o va lic ys ta h ia ta su sa di ni um s cr of oi de s n an no ce ra to ps is di ct ya m bo ni s ph al lo cy st a el on ga ta 123 210 205 200 195 190 185 180 t im e in m a r. rhaetica planorbis liasicus dstr r. rhaetica r. rhaetica dsj1 dsj2 dsj3 dsj4 dsj5 dsj6 dsj7 dsj8 dsj9 dsj10 p. nasuta p. nasuta l. spinosa l. spinosa l. spinosa m. semitabulatum n. gracilis m. reticulatum l. variabile l. variabile l. variabile d. priscum d. priscum d. priscum c hr on oz on es z on es ( th is s tu dy ) d sz on es z on es , d en m ar k su bz on es , d en m ar k z on es , g re at b ri ta in su bz on es , g re at b ri ta in lo ng -t er m e us ta tic c ur ve af te r h aq e t a l. (1 98 7) sh or tte rm e us ta tic c ur ve af te r h aq e t a l. (1 98 7) pa la eo te m pe ra tu re cu rv e (° c ) angulata levesquei thouarsence variabilis bifrons falciferum tenuicostatum spinatum margaritatus davoei ibex jamesoni raricostatum oxynotum obtusum turneri semicostatum bucklandi n. senex m. semitabulatum m. semitabulatum b a b b a a a a a b b b c se ale ve l c ur ve af te r h al la m ( 19 88 ) 20100 m100sea-level rise fig. 3. composite range chart for the key marker dinoflagellate cyst species in the lower jurassic of the subboreal region. in figs 3–5, the eustatic sea-level curves are from haq et al. (1987) and hallam (1988), the palaeotemperature curve is courtesy of b. buchardt (appendix 1) and the time-scale is from haq et al. (1987). note that danish jurassic ammonite recovery is sporadic and that only certain chronostratigraphic zones can be identified using macrofossils (poulsen 1996). is renamed the cribroperidinium crispum zone to accommodate the change in name of the index species (hedberg 1976; salvador 1994). the acanthaulax crispa (cribroperidinium crispum) zone was divided into subzones a and b by riding & thomas (1992) and the dsj13 and dsj14 zones correspond precisely to subzones a and b, respectively. dsj15–18; bathonian – middle callovian (zigzag–coronatum zones) zones dsj15, dsj16 and dsj17 correspond respectively to subzones a, b and c of the ctenidodinium sellwoodii zone of riding & thomas (1992; fig. 4). the latter biozone largely equates to the ctenidodinium combazii – 124 155 160 165 170 175 ju ra ss ic m id dl e ju ra ss ic a al en ia n ba jo ci an ba th on ia n c al lo vi an pe ri od ep oc h a ge t im e in m a m en di co di ni um r et icu la tu m n an no ce ra to ps is se ne x sc rin io ca ss is w eb er i m an co di ni um s em ita bu la tu m n an no ce ra to ps is gr ac ilis pa rv oc ys ta ‘s ui te ’ pa rv oc ys ta n as ut a ph al lo cy st a eu m ek es sc rin io ca ss is pr isc us o va lic ys ta h ia ta su sa di ni um s cr of oi de s n an no ce ra to ps is di ct ya m bo ni s ph al lo cy st a el on ga ta d ur ot rig ia d av ey i cr ib ro pe rid in iu m c ris pu m va lva eo di ni um s pi no su m li th od in ia c ay to ne ns is ‘g ro up ’ ct en id od in iu m c on tin uu m ca rp at ho di ni um p re da e al do rf ia a ld or fe ns is ct en id od in iu m c om ba zi i d ich ad og on ya ul ax s el lw oo di i ‘ gr ou p’ g on ya ul ac ys ta ju ra ss ica a de ct a n an no ce ra to ps is pe llu cid a im pl et os ph ae rid iu m v ar isp in os um ri ga ud el la a em ul a ct en id od in iu m o rn at um si rm io di ni um g ro ss i co m po sit os ph ae rid iu m p ol on icu m g on ya ul ac ys ta e ise na ck ii st ep ha ne lyt ro n sc ar bu rg he ns e g on ya ul ac ys ta c en tr ico nn at a li m bo di ni um a bs id at um w an ae a th ys an ot a tr ich od in iu m s ca rb ur gh en se sc rin io di ni um c ry st al lin um ctenidodinium sellwoodii zone of woollam & riding (1983). the dsj18 zone is coeval with the ctenidodinium continuum zone of riding & thomas (1992), which is broadly coincident with the ctenidodinium ornatum – ctenidodinium continuum zone of woollam & riding (1983). dsj19; upper callovian (athleta, lamberti zones) the dsj19 zone corresponds to the wanaea thysanota zone of riding & thomas (1992), formerly the wanaea thysanota zone (subzones a and b) of woollam & riding (1983). the subdivision of the wanaea thysanota zone, based on the range base of trichodinium scar125 155 160 165 170 175 t im e in m a c hr on oz on es z on es ( th is s tu dy ) d sz on es z on es , d en m ar k su bz on es , d en m ar k z on es , g re at b ri ta in su bz on es , g re at b ri ta in lo ng -t er m e us ta tic c ur ve af te r h aq e t a l. (1 98 7) sh or tte rm e us ta tic c ur ve af te r h aq e t a l. (1 98 7) pa la eo te m pe ra tu re c ur ve ( °c ) se ale ve l c ur ve af te r h al la m ( 19 88 ) zigzag discus dsj13 n. gracilis c. crispum n. gracilis a. crispa c. continuum w. thysanota w. thysanota lamberti athleta jason calloviense koenigi herveyi coronatum aspidoides hodsoni subcontractus progracilis tenuiplicatus morrisi parkinsoni garantiana subfurcatus humphriesianum sauzei loeviusculo discites dsj11 dsj12 dsj14 dsj15 dsj16 dsj17 dsj19 concavum murchisonae p. nasuta dsj10opalinum not zoned p. nasuta not zoned m. semitabulatum c d e a b a b c c. continuum dsj18 c. sellwoodii c. sellwoodii 20100 m100sea-level rise fig. 4. composite range chart for the key marker dinoflagellate cyst species in the middle jurassic of the subboreal region. burghensis, was discontinued by riding & thomas (1992). dsj20; lowermost oxfordian (mariae zone) the dsj20 zone equates to the wanaea fimbriata zone of woollam & riding (1983) and riding & thomas (1992). the definition of the upper boundary of the dsj20 zone is given below, in the text pertaining to the dsj21 and dsj22 zones. dsj21, dsj22; lower–middle oxfordian (cordatum, densiplicatum zones) the name of the liesbergia scarburghensis zone of riding & thomas (1992), originally the acanthaulax senta zone of woollam & riding (1983), is changed to the trichodinium scarburghensis zone due to the name change of the index species (hedberg 1976; salvador 1994). the inception of forms belonging to the systematophora areolata group was used to redefine the lower boundary of this zone by riding & thomas (1992). the distribution of the systematophora areolata group in the united kingdom and the danish embayment appears to be palaeoenvironmentally controlled (poulsen 1992, 1996; riding & thomas 1992) and this group is therefore excluded from the revised definition, given below. dsj21; lower oxfordian (cordatum zone) the base of this zone is defined by the range base of leptodinium subtile (fig. 5). the top of the dsj21 zone is defined by the range tops of gonyaulacysta centriconnata, limbodinium absidatum and wanaea thysanota and the inception of endoscrinium luridum. the age of the zone corresponds to the cordatum zone; poulsen (1996) presented a detailed discussion of the dinoflagellate cyst zonations at the middle–upper jurassic boundary. dsj22; middle oxfordian (densiplicatum zone) the base of this zone is defined by the last occurrences of gonyaulacysta centriconnata, limbodinium absidatum and wanaea thysanota and the range base of endoscrinium luridum (fig. 5). the top of zone dsj22 is defined by the range top of the lithodinia caytonensis group, and the inceptions of glossodinium dimorphum and scriniodinium inritibile. dsj23–27; middle oxfordian – lowermost kimmeridgian (tenuiserratum–baylei zones) the gonyaulacysta jurassica – scriniodinium crystallinum dinoflagellate cyst zone of woollam & riding (1983) is herein divided into five dsj zones. the zone and its three subzones were defined by woollam & riding (1983) and emended by riding & thomas (1988, 1992); the upper boundary of the zone was emended by poulsen (1991). by introducing further subdivision, a redefinition of the zone and its constituent subzones is therefore required here (fig. 5). dsj23; middle oxfordian (tenuiserratum zone) the base of zone dsj23 is defined by the range top of the lithodinia caytonensis group and the inceptions of glossodinium dimorphum and scriniodinium inritibile. the top of zone dsj23 is defined by the last occurrences of rigaudella aemula and trichodinium scarburghensis. zone dsj23 is coeval with the lsc (c) subzone of riding & thomas (1992). dsj24; upper oxfordian (glosense zone) the base of the dsj24 zone is defined by the apparent extinctions of rigaudella aemula and trichodinium scarburghensis; the top is defined by the range top of compositosphaeridium polonicum. it is equivalent to the scr (a) subzone of riding & thomas (1992). dsj25; upper oxfordian (serratum, regulare zones) the base of the dsj25 zone is defined by the range top of compositosphaeridium polonicum. the top is defined by the last occurrence of gonyaulacysta jurassica subsp. adecta and the first appearances of dingodinium tuberosum and occisucysta balia. the dsj25 zone is equivalent to the scr (b) subzone of riding & thomas (1992). dsj26; upper oxfordian (rosenkrantzi zone) the base of the dsj26 zone is defined by the range top of gonyaulacysta jurassica subsp. adecta and the first appearances of dingodinium tuberosum and occisucysta balia. the top of the zone is defined by the last occurrence of ctenidodinium ornatum and the inception of senoniasphaera jurassica. the dsj26 zone equates to the scr (c) subzone of riding & thomas (1992). 126 dsj27; lower kimmeridgian (baylei zone) the base of the dsj27 zone is defined by the range top of ctenidodinium ornatum and the inception of senoniasphaera jurassica. the top of the zone is defined by the last occurrences of gonyaulacysta eisenackii, nannoceratopsis pellucida and scriniodinium crystallinum, and the first appearances of cribroperidinium? longicorne and oligosphaeridium patulum. the dsj27 zone equates to the scr (d) subzone of riding & thomas (1992). dsj28, dsj29; kimmeridgian (cymodoce–autissiodorensis zones) a bipartite subdivision of the endoscrinium luridum zone (formerly the scriniodinium luridum zone) of woollam & riding (1983) was introduced by nøhrhansen (1986); the zone was subsequently expanded and emended by riding & thomas (1988) and poulsen (1991). the definition of the lower boundary of the endoscrinium luridum zone of riding & thomas (1992), and thus the dsj28 zone, is here extended to include the last occurrence of gonyaulacysta eisenackii. the definitions of the top of zone dsj28 (or the stephanelytron scarburghense subzone of nøhr-hansen 1986) and the base and top of zone dsj29 (or the perisseiasphaeridium pannosum subzone) follow poulsen (1991). the definition of the top of zone dsj29 is emended to include the first appearances of egmontodinium polyplacophorum and systematophora daveyi (fig. 5). dsj30–34; lower–middle volgian (elegans–fittoni zones) the dsj30–34 zones correspond respectively to subzones a to e of the glossodinium dimorphum (gdi) zone of riding & thomas (1992). these units are coeval with subzones b and c of the glossodinium dimorphum – dingodinium tuberosum zone of woollam & riding (1983). dsj35–37; middle volgian (albani–anguiformis zones) the dsj35, dsj36 and dsj37 zones are the equivalent of the dichadogonyaulax culmula and dingodinium? spinosum zones of davey (1979). the latter are also equivalent to the dichadogonyaulax? pannea zone of riding & thomas (1992; formerly the ctenidodinium culmulum – ctenidodinium panneum zone of woollam & riding 1983). in this zonation for the entire british– danish area, the subdivision of the dichadogonyaulax culmula and dingodinium? spinosum zones of davey (1979, 1982) is used, and the dichadogonyaulax culmula zone is further subdivided into two zones (dsj35 and dsj36; fig. 5). it should be noted that the boundary between the dichadogonyaulax culmula zone (dsj36) and the dingodinium? spinosum zone (dsj37) is herein placed at the top of the glaucolithus zone, rather than at the top of the okusensis zone, where this boundary was placed by poulsen (1991, 1992, 1996). dsj35; middle volgian (albani zone) the base of the dsj35 zone is defined by the range base of consistent dichadogonyaulax culmula and the apparent extinction of occisucysta balia. the zone is the equivalent of the dpa (a) subzone of riding & thomas (1992). dsj36; middle volgian (glaucolithus zone) the base of the dsj36 zone is defined by the youngest occurrences of leptodinium subtile and scriniodinium inritibile. dsj37; middle volgian (okusensis–anguiformis zones) the base of the dsj37 zone is defined by the range base of dingodinium? spinosum. within this zone, or at the top, the range top of senoniasphaera jurassica is observed. this apparent extinction appears to occur in older strata in the british area than in the danish onshore area. dsj38–dsk2;middle volgian – upper ryazanian (oppressus–icenii zones) the pareodinia dasyforma and gochteodinia villosa zones of davey (1979, 1982) were subdivided into three subzones. the names of the egmontodinium expiratum, rotosphaeropsis thula, and endoscrinium pharo subzones of davey (1979, 1982) have been changed to follow the international guidelines whereby zonal names are changed in order to conform with any changes in the name of the index species (hedberg 1976; salvador 1994). the broadly equivalent gochteodinia villosa zone of woollam & riding (1983) and riding & thomas (1992) 127 was divided into subzones a to c by woollam & riding (1983) and riding & thomas (1992). the boundaries between subzones a to c of riding & thomas (1992) are different to those of the egmontodinium expiratum, rotosphaeropsis thula and endoscrinium pharo subzones of davey (1979, 1982) and adopted by poulsen (1991, 1992, 1996). the top of the gochteodinia villosa zone equivalent is herein regarded as being coincident with the top of the icenii zone. this is one chronozone lower than in the subdivisions of davey (1979, 1982), woollam & riding (1983), riding & thomas (1992) and poulsen (1996). the oppressus to icenii zone interval is herein subdivided into a four-fold subdivision, the dsj38, dsj39, dsk1 and dsk2 zones (fig. 5). 128 130 135 140 145 150 li th od in ia c ay to ne ns is ‘g ro up ’ ct en id od in iu m c on tin uu m d ic ha do go ny au la x se llw oo di i ‘ gr ou p’ g on ya ul ac ys ta ju ra ss ic a ad ec ta n an no ce ra to ps is pe llu ci da ri ga ud el la a em ul a ct en id od in iu m o rn at um si rm io di ni um g ro ss i co m po sit os ph ae rid iu m p ol on ic um g on ya ul ac ys ta e ise na ck ii st ep ha ne lyt ro n sc ar bu rg he ns e g on ya ul ac ys ta c en tr ic on na ta li m bo di ni um a bs id at um w an ae a th ys an ot a tr ic ho di ni um s ca rb ur gh en se sc rin io di ni um c ry st al lin um w an ae a fim br ia ta g on ya ul ac ys ta ju ra ss ic a ju ra ss ic a le pt od in iu m s ub til e en do sc rin iu m lu rid um sc rin io di ni um in rit ib ile g lo ss od in iu m d im or ph um al do rf ia d ic ty ot a py ru m o cc isu cy st a ba lia d in go di ni um tu be ro su m se no ni as ph ae ra ju ra ss ic a cr ib ro pe rid in iu m ? lo ng ic or ne o lig os ph ae rid iu m p at ul um sy st em at op ho ra a re ol at a su bt ilis ph ae ra ? in af fe ct a su bt ilis ph ae ra ? pa em in os a pe ris se ia sp ha er id iu m p an no su m d ic ha do go ny au la x pa nn ea eg m on to di ni um p ol yp la co ph or um sy st em at op ho ra d av ey i pe ri od ep oc h a ge t im e in m a ju ra ss ic la te ju ra ss ic o xf or di an k im m er id gi an v ol gi an r ya za ni an v al . c re ta ce ou s ea rl y c re ta ce ou s eg m on to di ni um e xp ira tu m ro to sp ha er op sis th ul a m ud er on gi a sim pl ex kl ei th ria sp ha er id iu m p or os isp in um d ic ha do go ny au la x cu lm ul a d in go di ni um ? sp in os um g oc ht eo di ni a vil lo sa ba tio la di ni um r ad ic ul at um o cc isu cy st a sp . a o f d av ey 1 98 2 ps eu do ce ra tiu m p el lif er um dsj38; middle–upper volgian (oppressus, primitivus zones) the base of the dsj38 zone is defined by the youngest occurrences of dichadogonyaulax? pannea, dingodinium tuberosum and glossodinium dimorphum and the oldest consistent occurrence of gochteodinia villosa. this zone is therefore equivalent to subzone a of the gochteodinia villosa zone of riding & thomas (1992) and the lower part of the egmontodinium expiratum subzone of davey (1979, 1982). dsj39; upper volgian – lower ryazanian (preplicomphalus–runctoni zones) the base of the dsj39 zone is defined by the range top of egmontodinium polyplacophorum. this zone spans the jurassic–cretaceous boundary and is equiv129 lo ng -t er m e us ta tic c ur ve af te r h aq e t a l. (1 98 7) sh or tte rm e us ta tic c ur ve af te r h aq e t a l. (1 98 7) pa la eo te m pe ra tu re c ur ve ( °c ) se ale ve l c ur ve af te r h al la m ( 19 88 ) c hr on oz on es z on es ( th is s tu dy ) d sz on es z on es , d en m ar k su bz on es , d en m ar k z on es , g re at b ri ta in su bz on es , g re at b ri ta in t. scarburghense s. crystallinum s. crystallinum not zoned t. scarburghense e. luridum e. luridum g. dimorphum g. dimorphum d. culmula d. spinosum d. spinosum d. culmula g. villosa g. villosa g. villosa d. pannea g. dimorphum e. luridum s. crystallinum t. scarburghense s. scarburghense p. pannosum dsj37 dsj39 paratollia albidum stenumphalus icenii kochi runctonii lamplughi preplicomphalus primitivus oppressus anguiformis kerberus okusensis glaucolithus albani fittoni rotunda pallasioides pectinatus hudlestoni wheatleyensis scitulus elegans autissiodorensis eudoxus mutabilis cymodoce baylei rosenkrantzi regulare serratum glosense tenuiserratum densiplicatum cordatum mariae w. fimbriata w. fimbriatadsj20 dsj21 dsj22 dsj23 dsj24 dsj26 dsj27 dsj28 dsj29 dsj30 dsj31 dsj32 dsj33 dsj34 r. thula e. pharo dsj35 dsj36 dsj38 dsk1 dsk2 130 135 140 145 150 t im e in m a dsj25 e. expiratum not zoned d c b a b a a b b c d e a a a a a a b b b b b b c c c c c d d e 20100 m100sea-level rise fig. 5. composite range chart for the key marker dinoflagellate cyst species in the upper jurassic – lowermost cretaceous of the subboreal region. alent to the majority of subzone b of the gochteodinia villosa zone of riding & thomas (1992) and the upper part of the egmontodinium expiratum subzone and the lower rotosphaeropsis thula subzone of davey (1979, 1982). dsk1; lower ryazanian (kochi zone) the top of the dsk1 zone is defined by the youngest occurrence of rotosphaeropsis thula. this zone is equivalent to the upper part of subzone b of the gochteodinia villosa zone of riding & thomas (1992) and the upper part of the rotosphaeropsis thula subzone of davey (1979, 1982). dsk2; upper ryazanian (icenii zone) the base of the dsk2 zone is defined by the youngest occurrence of rotosphaeropsis thula and the oldest appearance of occisucysta sp. a of davey (1982). the top of this zone is defined by the oldest appearance of pseudoceratium pelliferum. the range top of systematophora daveyi is present at the lower boundary of this zone in the danish basin and the danish north sea. in the united kingdom, this species occurs in younger strata. the dsk2 zone is equivalent to the lower part of subzone c of the gochteodinia villosa zone of riding & thomas (1992) and the lower part of the endoscrinium pharo subzone of davey (1979, 1982). the jurassic dinoflagellate cyst zonation as proxy for palaeoenvironmental changes the causal background for the spatial and temporal distributions of fossil and modern biotas is a function of the interplay of many factors. among these parameters, salinity and water temperatures have direct importance in dinoflagellate cyst distribution patterns, whereas changes in sea level have an indirect influence by changing coastal to shelf environments (wall et al. 1977; stover et al. 1996). before considering the palaeoecological implications of jurassic dinoflagellate distributions, it is instructive to briefly outline the established palaeoclimatic and palaeo-oceanographic scenario for jurassic times in subboreal northwest europe. jurassic sea-level change jurassic sea-level changes are well-established in the literature and several high-resolution sea-level curves have been published (e.g. haq et al. 1987; hallam 1988, 1992b). furthermore, stratigraphic analyses of sedimentary basins via the recognition of genetically-related stratal packages bounded by unconformities (sequences) have recently been developed by many workers such as partington et al. (1993), andsbjerg & dybkjær (2003, this volume) and nielsen (2003, this volume). jurassic palaeoclimates and ammonite provincialism jurassic palaeoclimates were characterised by weaker temperature gradients and more uniform seawater temperatures than at present (berggren & hollister 1974; gordon 1975). furthermore, the polar regions were significantly warmer than today without indications of continental glaciation (arkell 1956; donn 1982; valdes & sellwood 1992). the numerical general circulation model presented by valdes & sellwood (1992) may indicate somewhat lower palaeotemperatures for the kimmeridgian than for other intervals of the jurassic. however, the latter estimates appear to be inconsistent with other measurements and the geological record (see below). organisms demanding a warm climate lived closer to the polar regions than today, even when continent migration due to plate tectonics is taken into account. for example, hermatypic corals have been reported from the bathonian of east greenland (håkansson et al. 1971), a remarkable record as the northern proto-atlantic ocean was isolated from the tethys ocean during bajocian and bathonian times (birkelund & perch-nielsen 1976; callomon 1985, 2003, this volume). stahl & jordan (1969) measured palaeotemperatures between 18°c and 24°c from isotopic studies of ammonites from the german aalenian and an average palaeotemperature of 13°c (with a range of 8–22°c) from studies of septa in a callovian ammonite from poland. tan et al. (1970) determined palaeotemperatures of 19–25°c from early callovian belemnites from staffin bay, skye, north-west scotland. these authors also demonstrated middle callovian palaeotemperatures of 21–24°c using belemnites, although ammonites from the same horizons revealed palaeotemperatures of 28–30°c; tan et al. (1970) noted, however, that the middle callovian ammonites are enriched in 13c, thereby giv130 131 ing slightly higher calculated palaeotemperatures. measurements of palaeotemperature based on material originating from palaeoenvironments with lower than fully marine palaeosalinities may give erroneously high values, some 2–3°c higher than normal (donn 1982). the subboreal sea during the jurassic exhibited significant palaeosalinity fluctuations (hallam 1969; fürsich & sykes 1977), which may explain the slightly higher measurements of tan et al. (1970). isotopic studies on coccoliths from the kimmeridge clay of westbury, wiltshire, england revealed a palaeotemperature of 20–30°c throughout the early kimmeridgian (salinas 1984). although the jurassic was characterised by less pronounced global marine temperature gradients than those of the present-day, palaeoecological studies and palaeothermometry have demonstrated that jurassic palaeoclimatic variations were nevertheless significant. the palaeotemperature curve for the jurassic in figure 6 has been compiled by b. buchardt (table 2, appendix 1). it suggests that the early jurassic was characterised by a general cooling until the latest pliensbachian, followed by a rise which terminated at the aalenian–bajocian boundary. there followed a rapid palaeotemperature fall during the bajocian, succeeded by minor fluctuations for the remainder of the middle jurassic. the late jurassic was characterised by a steady temperature rise until the middle volgian, after which time there was a period of cooling to the jurassic–cretaceous boundary (fig. 6; appendix 1). enay (1980) and hallam (1983) stated that jurassic palaeoclimates had a relatively minor influence on jurassic provincialism, the biotic endemism observed being largely controlled by isolation related to plate tectonic events and sea-level changes. similarly, fürsich & sykes (1977) found that factors such as regional palaeotemperature and/or palaeosalinity gradients cannot alone explain the existence of the boreal realm durpalaeotemperature (°c) 242220181614121086 hettangian sinemurian pliensbachian toarcian aalenian bajocian bathonian callovian oxfordian volgian kimmeridgian polewards direction equatorial cosmopolitan dinoflagellate cyst flora boreal and tethyan dinoflagellate cyst provincialism boreal, australian and tethyan dinoflagellate cyst provincialism cosmopolitan dinoflagellate cyst flora mediterranean and euro-caucasian ammonite provincialism cosmopolitian ammonite fauna reduced boreal ammonite diversity boreal ammonite provincialism tethyan ammonite impoverishment and boreal ammonite isolation subboreal ammonite province and migration tethyan ammonite migration boreal ammonite invasions boreal ammonite expansion tethyan ammonite migration tethyan ammonite migration palaeodepositional temperature la te ju ra ss ic m id dl e ju ra ss ic ea rl y ju ra ss ic fig. 6. jurassic palaeotemperature curve, courtesy of b. buchardt (appendix 1). the heavy black line connects the average isotopic temperature for each stage, whereas the grey envelope shows the scatter of individual isotopic results. true palaeotemperatures are believed to fall inside the grey envelope. ing the oxfordian. the late jurassic palaeotemperature rises, however, appear to have controlled the migration of tethyan faunas into the boreal realm. furthermore, falling palaeotemperatures during the jurassic caused boreal faunal expansions or invasions into the tethyan realm (see below). ammonites were cosmopolitan during the earliest jurassic (hettangian and sinemurian), but in the early pliensbachian and throughout the remainder of the jurassic, ammonite provincialism was well-developed (hallam 1971, 1973; enay 1972, 1980; callomon 1985, 2003, this volume; cariou et al. 1985; page 2003, this volume; zeiss 2003, this volume). the mediterranean and euro-caucasian provinces were developed during the pliensbachian and by the close of the stage a boreal fauna made its first southerly incursion into the tethyan realm. the establishment of ammonite provincialism was coincident with falling palaeotemperatures and the following ‘boreal expansion’ corresponds to the early jurassic temperature minimum (fig. 6). during the toarcian to early bajocian period, boreal ammonite faunas became less diverse as they expanded into the tethyan realm; this phenomenon is related to rising palaeotemperatures. furthermore, doming in the central north sea interrupted the passage from the boreal ocean to the tethyan area (sellwood & hallam 1974; whiteman et al. 1975; eynon 1981; ziegler 1988; underhill & partington 1993). during the bathonian, tethyan ammonite migration northwards towards the boreal ocean coincided with the opening of the passage through the united kingdom and the north sea, together with rising palaeotemperatures. the boreal cardioceratid ammonite fauna developed during the late bathonian and callovian and several boreal ammonite ‘expansions’ have been recorded. these early incursions coincide with falling palaeotemperatures at this time. in the latest callovian to kimmeridgian, palaeotemperatures rose and tethyan ammonite faunas migrated into the boreal realm. several boreal ammonite ‘expansions’ also occurred at this time. although the average palaeotemperature rose during this period, the maximum temperature fell during the latest callovian – late oxfordian (fig. 6); this probably explains biotic migration from both the boreal and tethyan realms. from the late oxfordian to the early middle volgian, the palaeotemperature rose and thereafter fell. during the kimmeridgian, palaeotemperatures rose and a subboreal ammonite fauna was developed which migrated towards both the boreal and tethyan realms. finally, during the volgian, falling palaeotemperatures resulted in isolation of the boreal realm, and tethyan impoverishment was recorded by the ammonite faunas (hallam 1971, 1973; enay 1972, 1980; sykes & callomon 1979; imlay 1980; callomon & birkelund 1982; birkelund & callomon 1985; callomon 1985, 2003, this volume; cariou et al. 1985; wierzbowski 1989; page 2003, this volume; zeiss 2003, this volume). jurassic dinoflagellate palaeoecology jurassic dinoflagellate cyst provincialism is almost negligible compared to ammonite endemism (enay 1972, 1980; davies & norris 1981). globally, jurassic dinoflagellate cyst assemblages are of broadly similar generic/specific composition and a large number of species are cosmopolitan. furthermore, they exhibit marked similarities in stratigraphic ranges throughout the world (riding & ioannides 1996). it appears that during the early jurassic, dinoflagellate cyst assemblages throughout northwest europe were broadly non-provincial. however, during the aalenian pre-rift central north sea (ziegler 1988; underhill & partington 1993), a land barrier blocked north sea marine communications and the dinoflagel132 stage average range number of references# temperature °c data points °c* volgian no data kimmeridgian 18 15–20 22 1, 2, 3, 5 oxfordian 16 13–18 18 1, 2 callovian+ 17 14–20 18 8 callovian‡ 8 6–18 13 2, 3, 4, 7 bathonian 12 8–17 8 1, 4, 9, 10 bajocian 15 10–18 20 1, 2, 3, 4, 6 aalenian 22 17–24 17 2, 3, 4, 7, 9, 10 toarcian 18 12–23 39 1, 2, 3, 4, 9, 10 pliensbachian 15 10–18 21 1, 2, 3, 4, 9, 10 sinemurian 17 13–21 6 2 hettangian 19 15–23 5 2 table 2. oxygen isotope palaeotemperatures for the jurassic of northwest europe * isotope palaeoptemperatures calculated from oxygen isotope values according to the equation given by craig (1965). + data from scotland, uk. ‡ data from germany. # references: 1: bowen (1961a, b) 2: fritz (1964) 3: jordan & stahl (1970) 4: kunz (1973) 5: salinas (1984) 6: spaeth et al. (1971) 7: stahl & jordan (1969) 8: tan et al. (1970) 9: veizer (1974) 10: veizer & fritz (1976) late cyst assemblages were consequently differentiated into the boreal and tethyan provinces (smelror 1993; riding & ioannides 1996). this barrier became submerged by rising sea levels during the callovian, and late callovian and oxfordian dinoflagellate cyst assemblages are cosmopolitan throughout europe and adjacent regions (raynaud 1978; smelror 1993; riding & ioannides 1996). however, during the kimmeridgian and volgian stages, tethyan and boreal–australasian floras became established (helby et al. 1987; riding & ioannides 1996). the assemblage diversity in the northern hemisphere also followed general ecological principles, with higher diversities in equatorial regions. for example, the early toarcian dinoflagellate cyst assemblages of southern germany appear to be more diverse than coeval associations from the united kingdom (riding 1987; riding & ioannides 1996). other palaeoecological factors, however, clearly controlled dinoflagellate cyst diversity, as the late triassic and the toarcian–aalenian floras from the sverdrup basin, arctic canada are both significantly more diverse than their european counterparts (davies 1983; riding & ioannides 1996). dinoflagellate cyst associations were affected by a number of inter-related factors such as latitude, climate (temperature), water depth, marine currents, nutrient supply, inter-basin seaways, barriers, distance from shoreline and salinity. for much of the jurassic, dinoflagellate cyst assemblages in the subboreal realm appear to have been influenced by these parameters, particularly palaeotemperature/climate, palaeobathymetry (sealevel fluctuations) and seaways. the distribution patterns of late jurassic dinoflagellate cysts with low and high surface relief are related to cold and warm water, respectively, according to dörhöfer (1977). lentin & williams (1980) established that high relief cyst surfaces and processes are a flotation adaptation necessary in warm water (the specific gravity of water is 1.00000 at 4°c, falling to 0.99567 at 30°c). this palaeoecological trend in the jurassic was noted by wierzbowski & århus (1990), smelror (1993) and riding & ioannides (1996); these authors found that complex process-bearing forms are more common in the middle–upper jurassic of the tethyan realm compared to the boreal realm. the seasonal variation in day length is another important ecological factor which is rarely considered in studies of phytoplankton provincialism (reid 1973). figure 2 illustrates the position of the pliensbachian, toarcian and volgian north poles, in addition to the approximate geographical extent of the pliensbachian and volgian two-month polar night, giving seasonal variations in day length and sunlight intensity. such variations will have increased towards the poles, independent of climate and climatic zones. this must have been an important factor in a period during which heat-demanding plants and animals lived at higher latitudes compared to the present. the seasonal variations in day length and sunlight intensity would have had a major influence on phytoplankton during jurassic times and thereby also on higher links in the food chain (hallam 1973; reid 1973). the position of the two-month polar night moved during the jurassic from a position in the jurassic arctic sea near north-east siberia to a position between siberia and alaska (fig. 2). this may have reduced biotic migration between the proto-arctic ocean and the pacific, especially in the latest jurassic, thereby compounding increasing provincialism in the latest jurassic caused by falling temperatures in the kimmeridgian–volgian (figs 2–6). biozonation, palaeotemperature and sea-level changes in the late sinemurian, eustatic fluctuations appear to have been critical for dinoflagellate cyst floras. at this time, a sea-level rise (fig. 3) appears to have controlled the earliest occurrence of liasidium variabile; this species is considered to be related to deeper marine conditions. its occurrence may be related to migration as a consequence of the rising sea-level. the latest sinemurian sea-level fall (fig. 3) may have caused the apparent extinction of liasidium variabile. cooling during the earliest jurassic may explain the apparent earlier extinction of dapcodinium priscum in the danish basin than in the united kingdom (see dsj1–2 zones, fig. 3). furthermore, dapcodinium priscum ranges up to the lowermost toarcian in portugal (davies 1985). dapcodinium priscum was apparently a temperature-sensitive species which was confined to a relatively narrow palaeotemperature window. thus it continued to live under warmer conditions in portugal, although it disappeared from the subboreal province in the early sinemurian due to palaeoenvironmental factors (riding & thomas 1992). the middle–late pliensbachian sea-level highstand corresponds to the inception of the genus nannoceratopsis, as manifested by the first appearance of nannoceratopsis senex. during the late pliensbachian – middle toarcian sea-level rise, several dinoflagellate cyst species appeared including other species of nannoceratopsis (dsj5–6 zones, fig. 3). several of these forms 133 have apparent extinctions in the earliest toarcian, possibly caused by a fall in sea level (fig. 3), and the latest pliensbachian palaeotemperature minimum (fig. 6). the lower toarcian (dsj7–8 zones) is marked by widespread indications of restricted marine conditions resulting from a rapid sea-level rise (gorin & feistburkhardt 1990). the succeeding middle toarcian to earliest aalenian short-term sea-level changes are reflected in the united kingdom by a greater degree of zonal subdivision (dsj8–10) than in the danish basin where zones dsj9 and dsj10 cannot be differentiated (figs 3, 4). the inception of the marker species nannoceratopsis dictyambonis, which defines the boundary between the dsj9 and dsj10 zones, and the biostratigraphically important species nannoceratopsis ambonis, scriniocassis priscus and scriniocassis weberi, are not observed in the danish basin (poulsen 1992, 1996), although they are characteristic elements both in the united kingdom (woollam & riding 1983; riding 1987) and germany (prauss 1989). sea-level change in more proximal (i.e. shallow-water) parts of the north sea region may have been an important controlling factor, especially in the danish basin, where ammonites and ostracods are also absent in the toarcian (sorgenfrei & buch 1964; michelsen 1975; poulsen 1996). furthermore, the palaeotemperature rises influenced changes in the carbon reservoir and may have had a controlling effect on the palaeoenvironment (gorin & feist-burkhardt 1990). this effect may especially have influenced the environment in denmark during the toarcian. palaeotemperatures rose weakly during the aalenian, fell sharply in the bajocian and weakly in the bathonian – middle callovian before rising again in the latest callovian (fig. 6). this is reflected in the northward migration of tethyan floras from the aalenian to the bajocian. this migration ceased later in the bajocian and bathonian (see below). the middle jurassic thermal doming in the north sea started in the late toarcian and led to non-marine to marginal marine sedimentation in the north sea area (underhill & partington 1993; nielsen 2003, this volume). the dinoflagellate cyst record is interrupted in denmark above the dsj10 zone. in the united kingdom, the zonation may be interpreted as an interaction between variations in palaeotemperature and sea level. the nannoceratopsis gracilis zone (dsj11–12 zones) broadly corresponds to the middle aalenian – early bajocian sea-level rise (fig. 4). the following sea-level maximum and fall (fig. 4) correlates to the dsj13–16 zones with the dsj14–15 zonal boundary corresponding to the middle bathonian shortterm sea-level minimum (fig. 4). the dsj17–20 zonal boundaries do not appear to have been related to known sea-level changes (fig. 4). fenton & fisher (1978) demonstrated that the bathonian transgressive phase controlled dinoflagellate cyst migrations from germany and southern england to eastern england and the central graben in the north sea. these authors also compared the aspidoides zone in the united kingdom, france, east greenland and spain and found that although certain species are consistently present, there are significant compositional differences in these assemblages. fenton & fisher (1978) noted that species of ctenidodinium in the united kingdom were more common south of the london– brabrant massif than to the north, with c. combazii and c. ornatum dominant, and that c. combazii was not recorded north of the massif. in the aquitane basin, south-west france, the assemblages comprise approximately 90% c. combazii, whereas to the north, in the paris basin near poitiers, the assemblages are dominated by acritarchs with only relatively rare representatives of ctenidodinium (including c. combazii) and lithodinia (valensi 1953; dupin 1965; fenton & fisher 1978). in eastern england, bathonian dinoflagellate cyst assemblages are dominated by species of chytroeisphaeridia, ctenidodinium, lithodinia and sentusidinium. fenton & fisher (1978) concluded that highly sculptured species were common in the southern region, whereas the boreal assemblages include many smooth forms and that the callovian transgression resulted in a northern spread of sculptured forms (fig. 7). investigations of the bathonian–callovian from northwest germany (gocht 1970; fenton 1981), the netherlands (herngreen & de boer 1978) and the united kingdom (sarjeant 1959, 1976; neves & selley 1975; lam & porter 1977; muir & sarjeant 1978; fenton et al. 1980; woollam 1980, 1982; riding 1982; riding et al. 1991) have identified significant differences between the tethyan ctenidodinium-dominated assemblages and their boreal counterparts. the latter include shallowwater, proximal assemblages with ctenidodinium continuum, c. ornatum, dichadogonyaulax sellwoodii, lithodinia spp., nannoceratopsis spp., pareodinia spp. and valensiella ovula. it appears that the tethyan floras migrated northwards from southern england during the latest bajocian – earliest bathonian. a mixed tethyan–boreal flora is recognised further north in oxfordshire and cambridgeshire, central england, in the upper bathonian – lower callovian and in scotland in the middle callovian. the northwards migration of ctenidodinium combazii and related species appears to have been in response to warming during the late 134 bajocian – middle bathonian (figs 6, 7). this northerly incursion of tethyan floras ceased in the middle bathonian and a mixed flora is present in central and northern england; this event coincides with falling palaeotemperatures as well as falling sea-level in the bajocian – middle callovian. the middle jurassic floras with common nannoceratopsis gracilis and valensiella ovula may be indicative of shallow-water, relatively nearshore settings, rather than these associations having boreal affinities. this assemblage type was displaced by a more diverse tethyan flora due to rising sea-levels in the bajocian (fig. 4). nannoceratopsis gracilis has been described as a euryhaline species, tolerant of reduced palaeosalinities (fisher 1980; bucefalo palliani & riding 1997). it was described by fensome (1979) as a dominant species in a restricted marine palaeoenvironment (see also surlyk et al. 1973). this taxon has been reported from shallow-water marine to deltaic swamp facies of yorkshire by hancock & fisher (1981). davey (1979), however, described n. gracilis as being abundant in normal marine sediments together with valensiella ovula and related forms. valensiella ovula was deemed to be a boreal species by norris (1975), although it has been reported from bulgaria (dodekova 1975) which was within the tethyan realm. furthermore, this species has not been recorded from northern north america. in east greenland, it was recorded by fensome (1979) in a shallow shelf palaeoenvironment (surlyk et al. 1981). palmer & jenkyns (1975) described the bathonian palaeoenvironment in the area north of the london– brabrant massif in england as fresh to brackish water lagoons (the oxfordshire shallows). towards the south, these lagoons became gradually more marine and were replaced by fully marine carbonate shelf palaeoenvironments in dorset, southern england. not only did the 135 500 km ? ? ? ? 40°n 0° 10°e10°w land paralic/deltaic siliciclastics coastline probable coastline speculative northward migration of tethyan dinoflagellate cyst flora middle callovian late bathonian middle bathonian early bathonian fig. 7. palaeogeographical map of the north sea area during the middle jurassic (modified slightly from callomon 2003, this volume) showing the progressive northwards migration of tethyan dinoflagellate cysts. facies change in relation to palaeotemperature variations, but also from non-marine to marginal marine and further to fully marine. the oxfordian and kimmeridgian stages were characterised by rising palaeotemperatures and relatively minor, short-term, sea-level changes (figs 5, 6). the dinoflagellate cyst zonation appears to be a reflection of both of these factors. provincialism was developed in the kimmeridgian–volgian interval. falling palaeotemperatures and repeated, short-term eustatic changes in the volgian appear to be expressed in the zonation. in the danish basin, lowermost middle oxfordian marine deposits succeed fluviatile middle jurassic deposits (poulsen 1992, 1996; nielsen 2003, this volume). at this time, the distribution of dinoflagellate cysts became uniform over the entire british–danish area (poulsen 1996). raynaud (1978) compared the callovian to volgian (tithonian) dinoflagellate cyst record in the united kingdom and the north sea and found broad similarities between the areas, although quantitative differences were discerned. these differences were most pronounced in the kimmeridgian to volgian (tithonian) interval. according to raynaud (1978), late callovian – early oxfordian floras were uniform throughout arctic canada, east greenland, europe and svalbard (beju 1971; johnson & hills 1973; tan & hills 1978; bjærke 1980; poulsen 1984). the rising palaeotemperatures during the late callovian and oxfordian, together with the marine connection between the boreal and tethyan oceans, appear to have created a cosmopolitan floral province during this interval in the northern hemisphere. the majority of upper jurassic dinoflagellate cyst zonal boundaries were probably influenced by late jurassic eustatic oscillations (fig. 5). for example, the dsj28–29 zonal boundary coincides with the short-term sea-level fall following the mid-kimmeridgian maximum. in contrast, the dsj29–30 zonal boundary (the top of the endoscrinium luridum zone) corresponds approximately to the sea-level maximum in the latest kimmeridgian. similarly, many of the boundaries of the dsj30–38 zones correlate broadly with sea-level minima or maxima of the short-term eustatic curve of haq et al. (1987; fig. 5). the upper boundary of the glossodinium dimorphum zone (the dsj34–35 zonal boundary) approximates to a short-term sea-level maximum. the boundary between the dingodinium? spinosum and the gochteodinia villosa zones (dsj37–38) coincides with the short-term sea-level maximum at the close of the volgian (fig. 5). most of the zonal boundaries that correlate broadly with eustatic events have been recorded slightly later than the sea-level minima and maxima, and may be related to shifts in sedimentary facies from clay to sand (e.g. from the kimmeridge clay to portland sand formations or from the børglum to frederikshavn formations in dorset and jylland, respectively). kimmeridgian – early volgian dinoflagellate cyst assemblages are rich; the high diversities reflect correspondingly high sea levels. rich nearshore assemblages of latest kimmeridgian – earliest volgian age have been recorded in denmark at the margin of the baltic shield (poulsen 1996), indicating the maximum extent of the børglum formation (kimmeridge clay formation equivalent) over the baltic shield. marine conditions existed continuously in denmark during the volgian to ryazanian, although non-marine phases occurred near the margin of the baltic shield (poulsen 1996). poulsen (1992; 1996) considered volgian and ryazanian palaeoenvironments and identified floral events in denmark which were thought to reflect eustatic changes. these events are more readily identified in shallow shelf deposits than in the deeper water succession of the central graben. low diversity dinoflagellate cyst assemblages in the lower part of the dsj39 zone were recorded in denmark by poulsen (1996), who related them to eustatic rise and fall during the latest jurassic. the rich assemblages within the upper part of the dsj39 zone correspond to the sea-level rise at the jurassic–cretaceous boundary (poulsen 1996). less diverse assemblages recorded from the lowermost cretaceous (dsk1–2 zones) are related to successive marine highstands and lowstands (poulsen 1996). the richer assemblages from the overlying pseudoceratium pelliferum subzone have been related to the succeeding sea-level rise (poulsen 1996). the lowermost cretaceous of denmark is characterised by rich dinoflagellate cyst assemblages which include a variety of pareodiniacean dinoflagellate cysts, thereby suggesting affinities to the boreal realm (poulsen 1996). this corresponds to falling palaeotemperatures close to the jurassic–cretaceous boundary (fig. 6). late ryazanian assemblages are characterised by a decrease in the abundance and diversity of dinoflagellate cysts in the dsk2 zone and younger strata. lott et al. (1989) also noted this phenomenon and related the differences to a latest ryazanian transgression. it is believed, therefore, that the dinoflagellate cyst diversity and abundance fluctuations correspond well to both the sealevel changes at the jurassic–cretaceous boundary proposed by haq et al. (1987), and to the model of rawson & riley (1982). 136 137 conclusions the stepwise evolution of dinoflagellate cyst assemblages, as defined by inceptions and apparent extinctions, appears to have been largely controlled by sea-level changes, particularly during intervals with significant short-term eustatic fluctuations. in times of less pronounced, or more long-term sea-level changes, fluctuations in the marine palaeotemperature seem to have influenced dinoflagellate evolution. differences in the ranges of certain taxa between denmark and the united kingdom are ascribed to minor palaeotemperature differences. the early jurassic was characterised by a general cooling until the late pliensbachian. the zonation can be related to sea-level changes, although the slightly different ranges of some of the index species in denmark compared to the united kingdom can be explained by minor palaeotemperature differences. the late toarcian – aalenian thermal doming event in the north sea resulted in non-marine to marginal marine sedimentation during late aalenian – bathonian times in the north sea and denmark, such that the dinoflagellate cyst zonation is often difficult to apply in this stratigraphic interval. in the united kingdom, however, the zonation can be interpreted to reflect the interaction between the general falling palaeotemperature and sea-level changes. the oxfordian–kimmeridgian was characterised by warming and short-term sea-level changes. the dinoflagellate cyst zonation appears to record both these factors, sometimes together, sometimes singly. the falling palaeotemperature and the short-term sea-level changes in the volgian are expressed in the zonation, in addition to increased provincialism. comparison of the zonation scheme with the shortterm sea-level curve demonstrates that the ds zones in the subboreal province (subzones in the earlier zonation scheme for the british–danish area, see figs 3–5) often correlate with the short-term sea-level minima of haq et al. (1987). however, many of the zonal boundaries in the earlier british–danish zonation scheme correlate with short-term sea-level maxima on the haq et al. (1987) sea-level curve. hallam (1983) stated that jurassic palaeoclimates had only a minor influence on jurassic provincialism, and that any biotic endemism was largely controlled by plate tectonic events and sea-level changes. however, increases in jurassic palaeotemperatures allied to palaeooceanographic factors appear to have controlled the migration of tethyan biotas northward into the boreal realm. furthermore, falling jurassic palaeotemperatures caused boreal spreads or invasions southwards into the tethyan realm. acknowledgements we wish to express our sincere thanks to colleagues at our respective institutes for their help and interest, and to the referees, roger j. davey and don g. benson, for their constructive reviews. j.b.r. publishes with the permission of the british geological survey. references including references cited in appendix 1 andrews, i.j., long, d., richards, p.c., thomson, a.r., brown, s., chesher, j.a. & mccormac, m. 1990: united kingdom offshore regional report: the geology of the moray firth, 106 pp. london: her majesty’s stationery office for the british geological survey. andsbjerg, j. & dybkjær, k. 2003: sequence stratigraphy of the jurassic of the danish central graben. in: ineson, j.r. & surlyk, f. 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(eds): geology of greenland, 304–339. copenhagen: geological survey of greenland. bjærke, t. 1980: mesozoic palynology of svalbard v. dinoflagellates from the agardhfjellet member (middle and upper jurassic) in spitsbergen. norsk polarinstitutt skrifter 172, 145–167. bowen, r. 1961a: oxygen isotope paleotemperature measurements on cretaceous belemnoidea from europe, india and japan. journal of paleontology 35, 1077–1084. bowen, r. 1961b: paleotemperature analyses of mesozoic belemnoidea from germany and poland. journal of geology 69, 75–83. brooks, j.r.v. & chesher, j.a. 1975: review of the offshore jurassic of the uk northern north sea. in: finstad, k.g. & selley, r.c. 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(in association with the danish modelling group) 1988: quantitative basin modelling in the danish central trough. section i, model description, concepts, input, and output parameters, 72 pp. unpublished confidential report (in archives of geological survey of denmark and greenland, copenhagen, denmark). zeiss, a. 2003: the upper jurassic of europe: its subdivision and correlation. in: ineson, j.r. & surlyk f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 75–114 (this volume). ziegler, p.a. 1988: post-hercynian plate reorganization in the tethys and arctic – north atlantic domains. in: manspeizer, w. (ed.): triassic–jurassic rifting; continental breakup and the origin of the atlantic ocean and passive margins. developments in geotectonics 22, 711–755. amsterdam: elsevier. manuscript received 9 june 1997; revision accepted 7 july 1999. 143 appendix 1: oxygen isotope palaeotemperatures from the jurassic in northwest europe (by bjørn buchardt) the oxygen isotope palaeotemperature curve presented in figure 6 of this paper was first constructed as part of a basin modelling project (b. buchardt in: yükler & speers 1988); note that the term palaeotemperature is used here to refer to the temperature registered by organisms living at the time of deposition (epstein et al. 1951; epstein & lowenstam 1953). the curve was compiled from previously published data, predominantly from older literature. the data reflect the state-of-the-art in isotope analyses in the 1960s and 1970s, at which time palaeotemperature studies were mainly based on macrofossils collected at outcrop. a brief description is given here of the principles behind, and the background of, the palaeotemperature curve. an oxygen isotope palaeotemperature is a numerical value calculated from the oxygen isotopic composition of a carbonate shell according to the equations formulated by epstein et al. (1951) and craig (1965) for calcite and horibe & oba (1972) for aragonite, and expressed in degrees celsius (°c). interpretation of this value as a ‘true’ palaeotemperature is limited by several factors such as: (1) vital effects of the actual organisms (shell carbonate precipitated out of equilibrium with the surrounding water), (2) seasonal-selective shell formation, (3) variations in the oxygen isotope composition of the ambient water, and (4) postdepositional alteration of the shell. of these factors, the last is deemed to be the most important. vital effects are well-known among large groups of calcareous-shelled organisms that should be avoided for palaeotemperature work. however, molluscs are known to exhibit vital effects to only a small degree (epstein et al. 1951, fritz & poplawski 1974, buchardt & weiner 1988) and are eminently suitable for palaeotemperature studies. among the molluscs, belemnites and ammonites are believed to have had a nektonic to nektobenthonic habitat, where seasonal variations in seawater temperatures probably were of less importance. they are therefore more suited for palaeotemperature work than, for example oysters, which lived closer to the coast, in waters probably affected by seasonality. calculation of isotopic palaeotemperatures presumes knowledge of isotopic compositions of both shell carbonate and ambient seawater. as the last parameter cannot be measured directly, indirect assessments are necessary. the world’s oceans can be viewed as a homogeneous reservoir with a constant oxygen isotope composition in space, but not in time. the amount of 18o-depleted water bound in glacial ice affects the isotopic composition of the oceans globally. oceans today have an oxygen isotope composition of 0‰ on the ä-scale relative to the smow standard (standard mean ocean water). during preglacial times, this value was lowered by at least 1‰ (shackleton & kennett 1975), and consequently an average seawater ä18o-value of -1‰ has been applied in the present calculations. this correction amounts to a calculated temperature difference of approximately 4°c (craig 1965). unfortunately, seawater oxygen isotope homogeneity is also affected by any mixing with freshwater from rivers and streams discharging into the sea. river water is normally depleted in 18o to a highly variable degree. each estuary system thus has its own characteristic salinity/ä18o relationship (mook 1968; israelson et al. 1994). during the jurassic period in northwest europe, palaeo-oceanographic conditions were complex, and local isotopic effects from river discharge were probably common (salinas 1984; nøhr-hansen 1986). however, by mainly using nektobenthonic organisms such as belemnites and ammonites, these effects should be minimised. isotopic palaeotemperatures calculated from analyses of belemnites have been claimed by several authors to be unreliable (longinelli 1969; spaeth et al. 1971). this critique has focused on possible postdepositional alteration of belemnite guards (stahl & jordan 1969; spaeth et al. 1971; veizer 1974). the controversy has never been resolved, and belemnite oxygen isotope palaeotemperatures should only be applied when the volume of belemnite data is large enough to allow cross-checks between different areas and lithologies. stahl & jordan (1969) pointed out that the metastable aragonite phase in fossils is more reliable for isotopic studies than calcite because recrystallization processes invariably lead to the modification of aragonite to calcite and therefore can easily be identified. well-preserved aragonite in fossil shells can thus be taken as indication of minimal recrystallization and therefore of an unmodified isotopic signal. therefore, aragonite from ammonites can be viewed as a more reliable source for oxygen isotope palaeotemperatures than calcite from belemnites. however, the number of jurassic ammonite analyses reported in the literature are few, and it is not possible to compile a palaeotemperature curve solely from published ammonite data. consequently, data from both belemnites and ammonites has been compiled, although belemnite data are excluded where differences in isotopic composition between ammonites and belemnites are significant. oxygen isotope data from jurassic fossils in northwest europe (e.g. britain, germany and poland) have been published by several authors (bowen 1961a, b; fritz 1964; longinelli 1969; stahl & jordan 1969; tan et al. 1970; jordan & stahl 1971; kunz 1973; tan & hudson 1974; veizer 1974; veizer & fritz 1976; salinas 1984). these reports form the database for the oxygen isotope palaeotemperature curve presented in figures 3–6 of this paper. only data from ammonites and belemnites have been included. in most cases, stratigraphic resolution of the data is possible to zonal level. the curve is constructed from average values (solid line) and entire ranges (shaded area) of isotopic results for each stratigraphic 144 level. in table 2, the oxygen isotope palaeotemperatures have been calculated for each stage as an average and a range. the isotope palaeotemperature curve as shown in figures 3–6 is one method of displaying an oxygen isotope dataset for a heterogeneous selection of ammonites and belemnites from the jurassic deposits of northwest europe. the fact that the calculated palaeotemperatures fall within a credible range (8–26°c), supports the validity of the data and the compilation of information from different geographical areas. the only major discrepancy in the dataset is in the callovian, where extremely low values (minimum 8°c) in germany and poland (jordan & stahl 1971; kunz 1973) contrast with results from scotland (maximum 25°c; tan et al. 1970). in this study, the scottish results are excluded because they probably reflect the influence of 18o-depleted freshwater in a large estuarine system (tan & hudson 1974). the curve demonstrates isotopic variations which probably represent cold seawater conditions during the pliensbachian and the bajocian–callovian in northwest europe. these cold intervals are in contrast to warmer seawater conditions during the sinemurian, the toarcian–aalenian and the oxfordian–kimmeridgian. the maximum palaeotemperatures occurred at the early–middle jurassic transition and during the early kimmeridgian. geological survey of denmark and greenland bulletin 28, 2013, 53-56 53 fingerprinting of corundum (ruby) from fiskenæsset, west greenland nynke keulen and per kalvig since the late 1960s, it has been known that pink and red corundum occur in the area near fiskenæsset (qeqertarsuatsiaat) in southern west greenland. corundum is hosted in the fiskenæsset complex, which is part of the archaean basement of the north atlantic craton. to date, c. 40 corundum localities with a wide range of quality are known in the area – a few localities yield stones of gem quality. the most promising locality, aappaluttoq, is likely to be mined in the foreseeable future by the canadian company true north gems (figs 1, 2a). red corundum of gem quality is called ruby; gem quality corundum of other colours (e.g. pink, yellow or blue) is called pink sapphire, yellow sapphire etc., while the blue gem corundum is sapphire. red, pink and blue corundum are also known in smaller quantities from other areas in greenland. the fiskenæsset complex the fiskenæsset complex (fig. 1) comprises a series of intrusive sheets of anorthosite, leucogabbro, gabbro and ultramafic rocks (myers 1985), and is interpreted as derived from a supra-subduction setting, while the associated amphibolites stem from a mid-oceanic ridge to island arc basalt precursor (polat et al. 2009). the greater fiskenæsset region was metamorphosed c. 2.85–2.80 ga ago at midto upper amphibolite-facies temperatures and pressures, reaching granulite facies conditions near the village of fiskenæsset (mcgregor & friend 1992; schumacher et al. 2011). at least one generation of the c. 2.71 ga felsic pegmatite sheets cuts the anorthosite, ultramafic rocks, amphibolite and gneisses and created reaction zones that developed aluminium-rich mineral assemblages derived from the aluminium in the anorthosite rocks (schumacher et al. 2011; fig. 2b). these reaction zone assemblages, associated with pegmatitic felsic sheets and the ultramafic bodies, include very coarse-grained, radial anthophyllite ± green pargasite ± green or red spinel ± sapphirine ± cordierite (up to 30 cm single crystals) ± pink corundum, and ± phlogopite (schumacher et al. 2011). this study is a first attempt to find geochemical and mineralogical characteristics that can be used to tie the greenlandic rubies to their area of origin. this may have practical implications if an operation of rubies and pink sapphires is established in greenland. here, we present laser ablation inductively coupled plasma mass spectrometry (la-icpms) trace-element geochemical and oxygen isotope data of samples from the fiskenæsset area and other known localities in greenland (storø, maniitsoq, kapisillit and nattivit). © 2013 geus. geological survey of denmark and greenland bulletin 28, 53–56. open access: www.geus.dk/publications/bull fig. 1. simplified geological map of the fiskenæsset area in southern west greenland showing the investigated pink and red corundum (ruby) localities in the fiskenæsset complex. map after keulen & kokfelt et al. (2011). m: maniitsoq, s: storø, k: kapisillit, f: fiskenæsset, n: nattivit. ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; ; 63°n63°n greenlandgreenland nnkkss mm ff 50°w50°w 10 km10 km qororssuaq eastqororssuaq east piqinikpiqinik rubin ø (tasiusarsuaq)rubin ø (tasiusarsuaq) qaqat aquleritqaqat aqulerit bjørnesund 2008bjørnesund 2008 qaqqatsiaqqaqqatsiaq siggartartuliksiggartartulikkigutilikkigutilik lower annertusoqlower annertusoq upper annertusoqupper annertusoq aappaluttoqaappaluttoq bjørn esun d bjørn esun d fiskenæsfjorden fiskenæsfjorden grædefjordgrædefjord gneissgneiss amphiboliteamphibolite granitegranite doleritedoleriteanorthosite (fiskenæsset complex)anorthosite (fiskenæsset complex) fiskenæssetfiskenæsset 5454 trace-element geochemistry corundum has the chemical formula al2o3 and like most other minerals usually includes very small quantities of other elements in its crystal structure. the amounts of these trace elements and their ratios may depend on the geological conditions during the formation of the corundum and therefore usually vary between individual corundum deposits. twenty-four different elements were analysed for by means of laicp-ms; however most of them were not detected, including sn, nb, and ta. our investigations of the greenlandic corundum were concentrated on the elements mg, si, ti, v, cr, fe, and ga, as these elements are present in significant amounts and are also the most widely documented. we used fig. 2. a: aappaluttoq, fiskenæsset, seen from a helicopter towards the north. the white-grey rock is anorthosite. b: rubies in their host rock at tasiussarsuaq, fiskenæsset, greenland. fig. 3. normalised trace-element distributions for a: ti-cr-ga, b: fe-cr-ga, c: fe-si-ga and d: fe-cr-ti in corundum from fiskenæsset. the data are compared with data on international and greenlandic corundum occurrences (calligaro et al. 1999; del castillo et al. 2009; kalvig & frei 2010; pornwilard et al. 2011; rakontondrazafy et al. 2008; schwarz et al. 2008; thirangoon 2008). different colours show different countries. initials of the authors’ names were used where more than one study of the same locality exist. diagrams were created with wxternary (keulen & heijboer 2011). a b aappaluttoq upper annertusoq lower annertusoq bjørnesund 2008 siggartartulik rubin ø kigutilik intex qororssuaq east qaqqatsiaq aappaluttoq (t) aappaluttoq (kf) annertusoq (kf) qaqat aqulerit (kf) bjørnesund 2008 (kf) rubin ø (kf) siggartartulik (kf) piqinik (kf) maniitsoq storø kapisillit nattivit winza umba tunduru sahabano zazafotsy jagdalek jaipur & mysore john saul tsavo ratnapura luc yen macedonia 80% fiskenæsset best aappaluttoq soamiakatra ilakaka andilamena pailin (p) pailin (c) bo rai chanthaburi mong hsu mogok ti* fe* cr* ga* cr* ga* si* fe* ga* 80 ti* fe* cr* 60 40 20 80 60 40 20 80 60 40 20 80 60 40 20 80 60 40 20 80 60 40 20 80 60 40 20 80 60 40 20 80 204060 80 20406080 204060 80 204060 a b c d 55 the laser ablation sector-field inductively coupled plasma mass spectrometer (la-sf-icp-ms) at the geological survey of denmark and greenland (frei & gerdes 2009), employing an element 2 instrument from thermo-fisher scientific and a up213 frequency-quintupled nd:yag solid state laser system from new wave research. data reduction and determination of concentrations were calculated off-line through the software iolite using the trace_elements_is routine (hellstrom et al. 2008). further details on the methods are found in keulen & kalvig (2013). results on the trace-element investigations of corundum grains separated from 21 hand specimens from ten localities in the fiskenæsset complex are shown with red symbols in the ternary diagrams of fig. 3. the data for corundum from the fiskenæsset complex are in good concordance with earlier data from the area (kalvig & frei 2010; thirangoon 2008). in fig. 3 they are compared with data from other localities in greenland and from internationally, well-known, ruby occurrences. samples from fiskenæsset show a considerably higher amount of cr (up to 14000 ppm) than samples from other areas in greenland and most international samples. the fiskenæsset rubies are relatively rich in fe and si, but relatively poor in ti and ga, while v and mg do not show very distinctive values compared to samples from other areas (kalvig & keulen, 2011). in order to use trace-element investigations as a fingerprinting tool for rubies it is necessary to investigate the amount of overlap between samples from fiskenæsset and other localities. the blue lines in fig. 3 include 80% (26 out of 32) of the samples from the fiskenæsset area, based on sample distribution density contouring. most samples from other localities plot outside the blue line, but an overlapping chemistry is found with samples from soamiakatra, ilakaka, and andilamena in madagascar, bo rai and chanthaburi in thailand, pailin in cambodia and winza in tanzania. rubies from all these localities are hosted by ultramafic to mafic rocks or are found as placer deposits. this indicates that the trace elements in the rubies derive from the ultramafic rocks that are associated with the anorthosite. however, if only the four handspecimens with the most transparent and most intensively red-coloured corundum grains from aappaluttoq, fiskenæsset, are taken into account, no overlap between these handspecimens and samples from other known ruby occurrences is seen. corundum from these handspecimens is closest in transparency and colour to the stones that would be sold from a potential mine and therefore represent the aappaluttoq signature. as these corundum grains have a distinct composition, it can be concluded that trace-element geochemistry with icp-ms is a helpful tool in fingerprinting rubies from greenland. oxygen isotope geochemistry oxygen isotopic composition measurements were performed on ten samples from greenland at the university of lausanne, switzerland using an isotope ratio mass spectrometer, employing a method similar to that described by kasemann et al. (2001), see kalvig & keulen (2011) for details. six of the samples come from the fiskenæsset complex. their δ18o values vary between 1.62 and 4.20‰ for the fiskenæsset area, which is low compared to the other areas in greenland (up to 10.03‰ for maniitsoq) with the exception of one sample from nattivit (2.41‰; fig. 4). the δ18o values are also low compared to most other investigated corundum deposits worldwide (giuliani et al. 2007). the lowermost values (δ18o <3‰) are nearly diagnostic for the fiskenæsset area – worldwide only the placer deposits at andilamena and ilakaka in madagascar and gem-corundum in a cordieritite from iankaroka, madagascar have lower reported δ18o values. low δ18o values (≤4‰) generally reflect rock types such as mafic rocks, mafic gneiss, basalts, and desilicated pegmatite in mafic rocks (giuliani et al. 2005), which is in excellent agreement with the mafic to ultramafic setting of the fiskenæsset rubies. the values for samples from nattivit, kapisillit and storø are also low (2.4, 4.5 and 6.0‰ respectively) and also plot in the mafic–ultramafic field. unfortunately, no further geological information is available for these specimens and the data can thus not be validated fig. 4. δ18o values for six samples from fiskenæsset and four other localities in greenland. the values are relative to vsmow (vienna standard mean ocean water). colours indicate the approximate colour of the stones. red boxes and classification as mafic-ultramafic, john saul mine, and marble after giuliani et al. (2007). mafic-ultramafic john saul mine aappaluttoq annertusoq bjørnesund 2008 rubin ø siggartartulik pikiniq nattivit maniitsoq kapisillit storø marble δ18o corundum (‰, v-smov) 497395 078667 513743 497393 497392 497383 497397 289933 224779 497396 0 6 16108 18 202 4 221412 24 ggu no. 5656 against field observations. the value for maniitsoq with δ18o =10.03‰ is typically related to skarns in marble, or to biotitite in gneiss related to shear zones with high fluid activity. the rubies in the investigated sample are assumed to stem from sapphirine-bearing hornblendite. the hornblendite was probably formed in a shear zone with high fluid activity (like the biotitites in madagascar). the low δ18o values are a potentially useful tool for fingerprinting greenlandic rubies, especially the very low values for the fiskenæsset complex and nattivit, as only few other international occurrences have such low values. conclusions high confidence fingerprinting of rubies requires a combination of independent analytical methods such as traceelement analyses, oxygen isotope analyses and other studies. the two methods discussed here are efficient in characterising the fiskenæsset rubies. the ongoing research focuses on optical and physical characteristics, spectroscopy methods and scanning xrf. acknowledgements this study is part of a collaboration project between the bureau of minerals and petroleum in nuuk, greenland and the geological survey of denmark and greenland. the authors wish to thank kerstin bauer and torsten vennemann for help with the oxygen isotope analyses. references calligaro, t., poirot, j.-p. & querré, g. 1999: trace element fingerprinting of jewellery rubies by external beam pixe. nuclear instruments and methods in physics research b 150, 628–634. calvo del castillo, h., deprez, n., dupuis, t., mathis, f., deneckere, a., vandenabeele, p., calderón, t. & strivay, d. 2009: towards the differentiation of non-treated and treated corundum minerals by ion-beaminduced luminescence and other complementary techniques. analytical and bioanalytical chemistry 394, 1043–1058. frei, d. & gerdes, a. 2009: precise and accurate in situ u-pb dating of zircon with high sample throughput by automated la-sf-icp-ms. chemical geology 261, 261–270. giuliani, g., fallick, a.e., garnier, v., france-lanord, c., ohnenstetter, d. & schwarz, d. 2005: oxygen isotope composition as a tracer for the origins of rubies and sapphires. geology 33, 249–252, http://dx.doi. org/10.1130/g21261.1. giuliani, g. et al. 2007: oxygen isotope systematics of gem corundum deposits in madagascar: relevance for their geological origin. mineralium deposita 42, 251–270. hellstrom, j., paton c., woodhead j. & hergt j. 2008: iolite: software for spatially resolved la(quad and mc) icpms analysis. in: sylvester p. (ed.): laser ablation icp–ms in the earth sciences: current practices and outstanding issues, 343–348. mineralogical association of canada short course series 40, 343–348. kalvig, p. & frei, d. 2010: testing the use of geochemical characteristics of corundum from greenland as a tool for geographical typing. danmarks og grønlands geologiske undersøgelse rapport 2010/68, 36 pp. kalvig, p. & keulen, n. 2011: aktiviteter i rubinprojektet 2011 – samarbejdsprojekt med råstofdirektoratet. danmarks og grønlands geologiske undersøgelse rapport 2011/138, 41 pp. kasemann, s., meixner, a., rocholl, a., vennemann, t., schmitt, a. & wiedenbeck m. 2001: boron and oxygen isotope composition of certified reference materials nist srm 610/612, and reference materials jb-2g and jr-2g. geostandards newsletter 25, 405–416. keulen, n. & heijboer, t. 2011: the provenance of garnet: semi-automatic plotting and classification of garnet compositions. geophysical research abstracts 13, egu 2011-4716. keulen, n. & kalvig, p. 2013: report of the activities in the ruby project 2012 – a joint project with the bureau of minerals and petroleum. danmarks og grønlands geologiske undersøgelse rapport 2013/09, 25 pp. keulen, n., kokfelt, t.f. & the homogenisation team 2011: a 1:100 000 seamless, digital, internet-based geological map of south-west and southern west greenland, 61°30´–64°n, http://geuskort.geus.dk/gisfarm/gis_svgreenland.jsp. copenhagen: geological survey of denmark and greenland. mcgregor, v.r. & friend, c.r.l. 1992: late archean prograde amphiboliteto granulite-facies relations in the fiskenæsesset region, southern west greenland. the journal of geology 100, 207–219. myers, j.s. 1985: stratigraphy and structure of the fiskenæsset complex, southern west greenland. bulletin grønlands geologiske undersøgelse 150, 72 pp. polat a., appel, p.w.u., fryer, b., windley, b., frei, r., samson, i.m. & huang, h. 2009: trace element systematics of the neoarchean fiskenæsset anorthosite complex and associated meta-volcanic rocks, sw greenland: evidence for a magmatic arc origin. precambrian research 175, 87–115. pornwilard, m.-m., hansawek, r., shiowatana, j. & siripinyanond, a. 2011: geographical origin classification of gem corundum using elemental fingerprint analysis by laser ablation inductively coupled plasma mass spectrometry. international journal of mass spectrometry 306, 57– 62. rakontondrazafy, a.f.m. et al. 2008: gem corundum deposits of madagascar: a review. ore geology reviews 34, 134–154. schumacher, j.c., van hinsberg, v.j. & keulen, n. 2011: metamorphism in supracrustal and ultramafic rocks in southern west greenland and south-west greenland 64°–61.5°n. danmark og grønlands geologiske undersøgelse rapport 2011/6, 29 pp. schwarz, d. et al. 2008: rubies and sapphires from winza, central tanzania. gems & gemology 44, 322–347. thirangoon, k. 2008: ruby and pink sapphire from aappaluttoq, greenland. status of on-going research. unpublished report for true north gems co. (in archives of geological survey of denmark and greenland, geus report file 23642, 18 pp.). authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ntk@geus.dk mailto:ksv@geus.dk geological survey of denmark and greenland bulletin 23, 2011, 29-32 29 interglacial deposits in denmark have traditionally been referred to the cromerian complex (hareskovian), holsteinian or eemian stages. however, based on studies of sediment cores from the deep sea many more than three quaternary interglacials have been documented, and in other parts of north-western europe it is becoming increasingly clear that the on-shore quaternary sequences are much more complex than previously believed. interglacial deposits are characterised by plant and animal remains indicating longer periods with climatic conditions similar to or warmer than today, whereas interstadial deposits were formed during shorter time spans and usually contain remains of relatively coldadapted, arctic or sub-arctic species. interglacial and interstadial deposits can be dated more or less precisely, and thus provide information about the relative age of glacial deposits. in 2010 the geological survey of denmark and greenland (geus) described samples from a 75 m deep borehole at 55°45.37´n, 12°19.63´e (elevation 20.7 m a.s.l.), at the address bakketoften 50 in måløv on north-eastern sjælland (fig. 1). coring was conducted for ballerup municipality (kommune) using the reverse circulation technique (‘omvendt skylning med lufthævning’) by the well-drilling company thomas brøker, and samples were collected every 2 m and sent to geus. a few shells of freshwater gastropods were noted, and since pre-holocene shell-bearing deposits are rare on sjælland, we decided to analyse the macrofossil content. we initially assumed that the sediments were deposited in a lake during a weichselian interstadial, because an interstadial deposit was reported from måløv by frederiksen & rosbirk (1999). however, we question this dating since frederiksen and rosbirk did not provide any data that could confirm an interstadial age or a lacustrine environment. this article gives details of macrofossil analyses of five samples from the borehole, which allow for more definite conclusions about the depositional environment of the site and its possible age. lithostratigraphy the drilling penetrated 52 m of quaternary sediments, 16 m of danian limestone and 7 m of cretaceous chalk (fig. 2). it stopped at a depth of 75 m below the ground surface. the quaternary sediments are dominated by glacial till and meltwater deposits. however, a unit of clay with a few shells of freshwater gastropods was found between 34 and 44 m. the clay unit is underlain and overlain by clayey till. material and methods five sediment samples were available for analysis of macrofossils, each weighing around 1 kg. the samples were soaked in a naoh solution at room temperature for two weeks and wet sieved on 0.4, 0.2 and 0.1 mm sieves. the residue left on the sieves was analysed using a dissecting microscope. the plant and animal remains studied are much larger than for example pollen grains or diatom frustules, and we call them macrofossils, even though for example ostracods are traditionally considered microfossils by palaeontologists. palaeoecology the results of the macrofossil analyses are presented in table 1. one of the samples did not contain any macrofossils, one contained only a few moss remains, one contained a new middle pleistocene interglacial sequence from måløv, sjælland, denmark ole bennike, esben lindgård, henrik jønsson granat, richard c. preece and finn viehberg 10°e 14°e 56°n 55°n 54°n 57°n sweden denmark germany50 km jylland måløv sjælland fyn førslevgaard freeport copenhagen fig. 1. map of denmark showing the location of måløv and other localities on sjælland with non-marine interglacial deposits discussed in the text. © geus, 2011. geological survey of denmark and greenland bulletin 23, 29–32. open access: www.geus.dk/publications/bull 3030 frequent moss remains and a few other macrofossils. the two most shallow samples were somewhat richer in macrofossils, but their concentration and diversity are low. the macrofossils are dominated by freshwater organisms, but at least one brackish-water species, the ostracod cyprideis torosa, is also present. it occurred with articulated carapaces, suggesting that reworking is unlikely. the carapaces and shells of c. torosa are noded (forma torosa), a feature indicating a salinity <7‰ (meisch 2000; frenzel et al. 2010). shells of freshwater gastropods are also present. some are fragmented, but this damage may have occurred during coring. the presence of both brackish-water and freshwater organisms implies low salinity conditions that would allow such a co-occurrence. a similar situation is found in many places today, for example at river mouths, in estuaries and in the gulf of bothnia. the lack of head capsules of chironomidae, as well as carapaces and head shields of cladocera, may also imply weakly brackish waters. interglacial lake deposits usually contain hundreds of head capsules of non-biting midge larvae (chironomids) per millilitre sediment, and thousands of cladoceran remains per millilitre. although it is risky to use negative evidence, we suggest that the absence of chironomid and cladoceran remains is due to brackish-water conditions. on the other hand, the lack of these remains could perhaps also be attributed to poor preservation, but the other macrofossils are well preserved, and this possibility is considered unlikely. no rivers are found on sjælland today, and it may be speculated if the deposit formed near the outlet from a former freshwater or brackish water baltic sea, somewhat similar to the palaeogeographical situation in the early holocene. the assemblage of freshwater taxa comprises at least 12 species. macrolimnophytes are represented by a species of nymphaeaceae (water lily, 1 leaf hair) and stratiotes aloides (water soldier, 3 leaf-margin spines). these water plants grow in shallow water, in lakes or in streams with slowly flowing water. both are typical of mesotrophic to eutrophic waters. freshwater ostracods are represented by darwinula stevensoni and scottia tumida (fig. 3) and freshwater molluscs by the gastropod borysthenia naticina (fig. 4) and the bivalves d ep th b el o w t er ra in s u rf ac e (m ) clayey till sandy till clayey till sandy till sand meltwater gravel clayey till clay clayey till meltwater sand danian limestone cretaceous chalk 0 10 20 30 40 50 60 70 a n al ys ed s ec ti o n fig. 2. lithological log of core dgu 200.5351 from måløv. sample no. 38775 38776 38777 38778 38779 depth (m) 34–36 36–38 38–40 40–42 42–44 mosses bryum sp. – – – r c polytrichum s.l. sp. 1 – – – – sphagnum sp. 2 3 – – – vascular plants nymphaeaceae indet. 1 – – – – potentilla sp. – 1 – – – rumex maritimus – 4 – – – stratiotes aloides 2 – – – 1 typha sp. 2 – – – – juncus sp. 1 – – – – ostracods cyprideis torosa 4 – – – – scottia tumida r 1 – – – darwinula stevensonia 1 – – – – molluscs borysthenia naticina 9 3 – – – sphaerium cf. solidum 1 – – – – sphaerium cf. corneum r – – – – sphaerium sp. – r – – – pisidium supinum 2 – – – – pisidium moitessierianum 2 – – – – bivalvia indet. – r – – – bryozoans plumatella repens 2 1 – – 3 fredericella indica – – – – 2 cristatella mucedo – 1 – – – table 1. macrofossils in five samples from interglacial deposits near måløv r: rare, c: common. 31 sphaerium cf. solidum, sphaerium cf. corneum, pisidium supinum and pisidium moitessierianum. these are species characteristic of fluvial environments. statoblasts of three species of bryozoans were found: cristatella mucedo, plumatella sp. and fredericella indica. the bryozoans may have lived on water plants. fredericella indica is rarely recorded as a fossil, probably because its statoblasts are indistinctive. however, it has been found in middle weichselian interstadial deposits in sweden, in late-glacial deposits in norway and in holocene deposits in norway, denmark and greenland. we cannot say if the remains of freshwater plants come from plants that grew in the depositional basin or if they were washed into the basin from freshwater environments in the catchment area. it is also possible that some of the remains of invertebrates, notably the statoblasts, may have been transported from lakes or streams into the basin. most of the freshwater species can occur in many different freshwater biotopes, but overall they are characteristic of a low-energy fluvial environment. the presence of c. torosa suggests some influence by marine water. the fine texture of the sediment (clay) implies deposition in a low-energy environment, such as a river with weak bottom currents. the presence of clay rather than gyttja indicates deposition in a river rather than in a lake. non-aquatic moss taxa are dominated by the mosses bryum sp. and sphagnum sp. but a leaf of the moss polytrichum s.l. sp. was also recovered. vascular land plants are represented by one achene of potentilla sp., one seed of juncus sp., two fruits of typha sp. and four fruits of rumex maritimus. the latter may indicate salt-water influence, supporting the evidence from c. torosa, although it is also found at inland sites today. several of these species probably grew in mires along the shore of the former water body where the clay was deposited. in particular, sphagnum sp. and typha sp. are characteristic of mires. the assemblage from måløv is remarkably similar to that of the sidestrand hall member of the cromer forest-bed formation in britain that has also yielded stratiotes aloides, borysthenia naticina, sphaerium solidum, cyprideis torosa and scottia (preece et al. 2009). the depositional environment and climate must have been extremely similar but these two deposits are not necessarily of the same age. palaeoclimate several of the species recovered are warmth-demanding, especially s. aloides, typha sp. and r. maritimus. they are widespread in denmark today and are also found in the southern and eastern parts of sweden. the mean july temperature at the northern range limit of pisidium supinum is around 15°c. b. naticina no longer lives in denmark but has a wide modern range in central and eastern europe, from southern and eastern germany and poland to hungary, rumania and south-western european russia (zilch & jaeckel 1962). the presence of this species indicates a more continental climate with summer temperatures higher than those in denmark today. c. torosa is also a warmth-demanding species (frenzel et al. 2010). age estimate the assemblage recovered from måløv clearly indicates deposition during an interglacial rather than interstadial period as was previously suggested for deposits at måløv (frederiksen & rosbirk 1999; probably the same deposit). several of the thermophilous species present, such as borysthenia naticina, sphaerium cf. solidum, stratiotes aloides, typha sp. and rumex maritimus, are unknown from interstadial contexts. interstadial deposits so far described from eastern denmark are characterised by arctic species, such as salix polaris, dryas octopetala and betula nana (bennike et al. 1994, 2007), none of which were found at måløv. fruits of stratiotes spp. are well known from interglacial deposits in europe, and leaf-margin spines have also been reported (bennike & hoek 1999). in denmark, fruits of s. aloides have been reported from the last interglacial, the eemian and from reworked pleistocene floras (hartz 1909). 250 μm fig. 3. scanning electron microscope images of ostracode shells from måløv. a: scottia tumida (internal view). b: scottia tumida (external view). c: cyprideis torosa (external view, juvenile). 2 mm fig. 4. light photographs of two shells of borysthenia naticina from måløv. 3232 outside denmark the species is known from the last interglacial and from several older interglacials. biostratigraphically, borysthenia naticina (fig. 4) is the most important species. this species has not previously been reported from denmark but is known as a pliocene fossil from southern russia, from the early pleistocene of northern france and the netherlands and from the middle pleistocene of central and eastern europe. there are middle pleistocene records to the north-west of its modern range in the rhine valley. in britain it has been reported from deposits referred to the cromerian complex (marine isotope stages 15 and 13), the hoxnian (marine isotope stage 11) and to marine isotope stage 9 (roe et al. 2009). in the netherlands, its youngest occurrence is in a deposit referred to marine isotope stage 7 (meijer 2003). in north-western europe, b. naticina is thus unknown from the last interglacial stage, the eemian, which strongly suggests that the deposit at måløv is of preeemian age. scottia tumida is an extinct species and so far only known from pleistocene interglacial deposits from germany, poland, the united kingdom, hungary and greece (kempf 1971). other interglacial non-marine deposits on sjælland interglacial deposits are quite common in denmark, but most of them are found in the western and southern parts of the country where erosion by advancing glaciers and meltwater was less intense than in the eastern parts. from sjælland, only a few interglacial deposits have been reported (fig. 1). most of them are marine deposits that have been referred to the eemian or holsteinian. two interglacial lake deposits have been reported from sjælland, at the copenhagen free port and at førslevgaard on southern sjælland (fig. 1). the mollusc fauna from these deposits includes the bivalve corbicula fluminalis (hartz 1909), which indicates an early or middle pleistocene age (meijer & preece 2000). conclusions we conclude that the clayey deposit found at måløv was deposited in a low-energy fluvial environment influenced by weakly brackish water. the occurrence of the gastropod borysthenia naticina suggests that the deposit is at least middle pleistocene in age. the mean july temperature was higher than in denmark today, and the deposit is clearly interglacial rather than interstadial. references bennike, o. & hoek, w. 1999: late-glacial and early holocene records of stratiotes aloides l. from north-western europe. review of palaeobotany and palynology 107, 259–263. bennike, o., houmark-nielsen, m., böcher, j. & heiberg, e.o. 1994: a multi-disciplinary macrofossil study of middle weichselian sediments at kobbelgård, møn, denmark. palaeogeography, palaeoclimatology, palaeoecology 111, 1–15. bennike, o., houmark-nielsen, m. & wiberg-larsen, p. 2007: a middle weichselian interstadial lake deposit on sejerø, denmark: macrofossil studies and dating. journal of quaternary science 22, 647–651. fredriksen, j. & rosbirk, e. 1999: fundering af bro i interstadialt søbassin. varv 1999(2), 59–63. frenzel, p., keyser, d. & viehberg, f.a. 2010: an illustrated key and (palaeo) ecological primer for postglacial to recent ostracoda (crustacea) of the baltic sea. boreas 39, 567–575. hartz, n. 1909: bidrag til danmarks tertiære og diluviale flora. danmarks geologiske undersøgelse ii. række 20, 292 pp. kempf, e.k. 1971: ökologie, taxonomie und verbreitung der nichtmarinen ostrakoden-gattung scottia im quartär von europa. eiszeitalter und gegenwart 22, 43–63. meijer, t. 2003: the late middle pleistocene non-marine molluscan fauna of borehole noorderhoeve-19e117 (province of noord-holland, the netherlands). cainozoic research 2, 129–134. meijer, t. & preece, r.c. 2000: a review of the occurrence of corbicula in the pleistocene of north-west europe. geologie en mijnbouw / netherlands journal of geosciences 79, 241–255. meisch, c. 2000: freshwater ostracoda of western and central europe. in: schwoerbel, j. & zwick, p. (eds): süßwasserfauna von mitteleuropa 8(3), 522 pp. heidelberg: spektrum akademischer verlag. preece, r.c., parfitt, s.a., coope, g.r., penkman, k.e.h., ponel, p. & whittaker, j.e. 2009: biostratigraphic and aminostratigraphic constraints on the age of the middle pleistocene glacial succession in north norfolk, uk. journal of quaternary science 24, 557–580. roe, h.m., coope, g.r., devoy, r.j.n., harrison, c.j.o., penkman, k.e.h., preece, r.c. & schreve, d.c. 2009: differentiation of mis 9 and mis 11 in the continental record: vegetational, faunal, aminostratigraphic and sea-level evidence from coastal sites in essex, uk. quaternary science reviews 28, 2342–2373. zilch, a. & jaeckel, s.g.a. 1962: mollusca. die tierwelt mitteleuropas 2(1), 294 pp. leipzig: verlag von quelle & meyer. authors’ addresses o.b., e.l. & h.j. g., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: obe@geus.dk r.c.p., department of zoology, university of cambridge, downing st., cambridge cb2 3ej, uk. f.v., institute of geology and mineralogy, university of cologne, zülpicher str. 49, 50674 cologne, germany. geological survey of denmark and greenland bulletin 7, 2004, p 49-52 49 pollen records from lake sediments have a great potential for providing information on the quantitative composition of past vegetation and land cover in the surrounding landscape. this can contribute to a better understanding of the development of the cultural landscape and interactions between human impact on the landscape and natural conditions like soil and climate. a good understanding of the history of cultural landscapes is necessary for choosing appropriate management strategies for areas dependent on cultural impact, such as heaths, meadows and dry pastures. it is also important for archaeological research concerning utilisation of the landscape in earlier periods. furthermore, quantitative reconstructions are relevant for climate research. here they can be used to test climate models, since model predictions of past climate can be translated into past vegetation, which can then be compared to pollen-based reconstructions. past vegetation cover is also a necessary input to climate models, as it influences albedo, evapotranspiration and carbon storage and cycling. quantifying vegetation from fossil pollen samples requires a detailed understanding of the way vegetation is reflected in pollen assemblages, including the approximate size of the area of vegetation represented. the relationship between pollen and vegetation is complicated by the fact that different plant species produce different amounts of pollen, and that pollen types are dispersed differently in the atmosphere, depending on their size, shape and weight. these pressing challenges in pollen analysis have attracted much attention in recent years. models have been developed to describe and simulate species specific pollen dispersal, to quantitatively relate pollen proportions to plant abundance, as well as estimate pollen productivity and to quantify the pollen source area of different types of basins (parsons & prentice 1981; prentice & parsons 1983; prentice 1985; sugita 1993, 1994; sugita et al. 1997, 1999; broström 2002; bunting et al. 2004). the geological survey of denmark and greenland (geus) has in recent years contributed to the development and validation of such models through the project agrar 2000 (odgaard 1999; nielsen 2003), where quantitative estimates of past land cover in different regions of denmark were one of the main objectives, and through participation in the international research network pollandcal (pollen landscape calibration), funded by nordforsk (nordic research board), which focuses on further model development, validation and application. the historical analogues approach modern analogues of past landscapes are very rare in denmark due to extensive changes to the cultural landscape during the past two centuries. these include intensification of agricultural practices, such as drainage, fertilisation and use of pesticides, plantations of introduced tree species and reductions in areas of pasture, wetland and heathland. the resulting changes in species composition, vegetation structure and growing conditions have affected the relationship between vegetation and pollen sedimentation through changes in pollen productivity and dispersal (nielsen & odgaard 2004). quantifying the relationship between pollen sedimentation in lakes and land cover using historical maps anne birgitte nielsen geological survey of denmark and greenland bulletin 7, 49–52 (2005) © geus, 2005 fig. 1. map showing the location of the 30 danish calibration sites and nine test sites used to assess the use of historical land cover and pollen assemblage data and the erv model for vegetation reconstruction. modified from nielsen & odgaard (2005). 50 km 12°e10°e8°e 56°n 55°n 57°n denmark swedentest sites calibration sites navnsø store økssø skånsø kragsø dallerup sø gudme sø avnsø store gribsø gundsømagle sø modern pollen datasets are thus not suitable as a standard for reconstructing cultural landscapes from fossil pollen samples. these problems may be overcome by using historical data (odgaard & rasmussen 2000; nielsen & odgaard 2004), although this approach involves several sources of error and potential bias, including a lack of spatial precision and detail, problems in interpretation of land cover signatures, biased species selection, changes in nomenclature and potential dating problems. in spite of these problems, most of which arise because the historical data were usually not created with the purpose of describing the plant species distribution in detail, this and other studies have shown that historical analogues can provide new insights and assist the interpretation of palaeoecological data. the historical calibration dataset collected for this study consists of pollen samples from around a.d. 1800 (identified by 210pb dating) from 30 small (3–30 ha) danish lakes (fig. 1), and land cover data from historical maps around the lakes (nielsen 2003). the historical maps used were the so-called ‘parish maps’, which show areas of different land cover, such as arable fields, forest, meadow and heath around a.d. 1800 at scale 1:20 000. spatial plant distribution was estimated from the land cover using historical data (b. fritzbøger and j.r. rømer, personal communication 2000) and comparison to modern analogues of old cultural landscape types in southern sweden (broström 2002), where areas of traditional land use are more common than in denmark. the nine study sites of the agrar 2000 project (nielsen & odgaard 2005) were not included in the calibration dataset, so these could be used as independent test sites (fig. 1) for the vegetation reconstructions. distance weighting vegetation data plants growing near a depositional basin contribute more to the pollen assemblages in the basin than plants farther away. vegetation data should therefore be distance weighted to reflect the ‘pollen sample’s view’ of the landscape, before pollen/vegetation calibration is applied (prentice 1985). the vegetation data from the historical maps were distance weighted according to the prentice/sugita model of pollen dispersal and deposition (prentice 1985; sugita 1993) using computer programs developed by s. sugita (1994, unpublished data 2002). in this model, pollen dispersal is assumed to follow sutton’s (1953) equation for the dispersal of small particles in the atmosphere, and depends on the weight and size of pollen grains, which are species specific, and on wind speed and atmospheric conditions. only windborne pollen is considered by the model, so only lakes with small or no inlet streams were used as calibration sites. the model's predictions of pollen assemblages deposited in different basins have been validated for forested landscapes in north america (calcote 1995; sugita et al. 1997; davis 2000) and for cultural landscapes in sweden and denmark (broström 2002; nielsen 2004; nielsen & sugita in press). the extended r-value model the quantitative relationship between the pollen assemblages from the lakes and the distance weighted plant abundance estimated from the historical maps was analysed using the extended r-value (erv) model (parsons & prentice 1981; prentice & parsons 1983; sugita 1994). this model was developed to achieve quantitative vegetation reconstructions from pollen proportions. the basic assumption is that the pollen loading of species k at site i (ρik) is linearly related to the distance weighted plant abundance of species k in the relevant source area of site i (xik): ρik = αixik + yio (1) where αi is the pollen productivity of species i, and yio is the amount of pollen of species i originating outside the relevant source area. both are assumed to be constant among sites in a region. if yio is furthermore assumed to form a constant proportion (zi) of the total plant abundance at the site (erv submodel 2), it is possible to relate pollen percentage to vegetation percentage, and estimate αi and zi from a dataset of pollen counts and distance weighted plant abundance, using a maximum likelihood method (prentice & parsons 1983). once these parameters are estimated, the inverse form of the erv model can be used to reconstruct plant proportions from pollen proportions. pollen productivity (αi) and the background component (zi) for four groups of plants (trees, poaceae, cerealia and calluna) were estimated from the a.d. 1800 vegetation and pollen data (table 1; nielsen & odgaard 2005). the estimated values of αi are relative, as one taxon (here poaceae) is chosen as reference taxon. the relative pollen productivity estimates of cerealia and calluna are lower than those for southern sweden (broström et al. 2004), whereas the estimate for trees in denmark is higher than that for sweden (sugita et al. 1999). the differences may be explained by the swedish estimates being based on analyses of moss polsters, while the danish estimates derive from lake sediments. tree pollen may be more easily dispersed to lakes than herb pollen, because it is released at a greater height. a difference in species composition within the plant groups between denmark in a.d. 1800 and sweden today may also contribute to the observed differences (nielsen 2004). 50 estimating pollen source area the spatial scale reflected by pollen samples is vital for quantitative reconstructions of past vegetation and landscape inferred from fossil pollen data, and a rigorous definition of the pollen source area is needed. sugita (1994) proposed the concept of the ‘relevant source area of pollen’ (rsap), defined as the area beyond which the correlation between pollen deposition at each site and the surrounding vegetation does not improve. the pollen loading (in terms of amount and composition) coming from beyond rsap is constant between sites within a region, corresponding to yio in equation (1). the radius of rsap can be estimated simultaneously with erv model parameters from the dataset of pollen counts and distance weighted plant abundance, and the distance where the likelihood function score of the erv calculations no longer decreases, because fit of the model no longer improves (sugita 1994). by analysing the data from the calibration sites, the radius of rsap for the danish lakes is estimated to c. 1800 m from the centre of lakes for all sites (nielsen & sugita in press). a difference in rsap radius of 400–500 m between eastern and western denmark was observed. this can be explained by a difference in the average patch size of the vegetation between regions (33 ha around eastern sites, 79 ha around western sites). both simulations (sugita 1994; broström 2002; bunting et al. 2004; nielsen & sugita in press) and empirical data (calcote 1995; nielsen & sugita in press) suggest that rsap is affected primarily by the spatial distribution of vegetation, especially patch size. rsap has been shown to be largely independent of fall speed or relative pollen productivity of the taxa present in the landscape, so the species involved have little effect on rsap (bunting et al. 2004; nielsen & sugita in press). the size of rsap is important for the interpretation of fossil pollen assemblages, as it is only vegetation within this area that potentially can be reconstructed. reconstructing land cover the estimates of αi and zi from the calibration sites were used to reconstruct vegetation composition around the nine test sites, applying the inverse form of the erv model (prentice & parsons 1983). the reconstructions are compared to distance weighted vegetation composition in rsap of the test sites from historical maps (fig. 2). the reconstructions based on the erv model reflect actual differences in vegetation among the test sites much more clearly than the pollen pro51 αi (sd) zi (sd) α, southern sweden poaceae 1.0 (0.0) 0.94 (0.022) 1.0 cerealia 0.95 (0.20) 0.11 (0.025) 3.2 calluna 2.06 (0.042) 0.10 (0.005) 4.7 trees 9.41 (0.48) 1.33 (0.002) 5.95 pollen productivity (αi) and background component (zi) estimated from the calibration sites, using erv submodel 2. for comparison, pollen productivity estimated from moss samples from southern sweden (sugita et al. 1999; broström et al. 2004) are also listed. table 1. erv parameter estimates fig. 2. a: pollen proportions in the a.d. 1800 sediment of the nine test sites. b: distance weighted vegetation proportions reconstructed from the pollen counts using the reverse erv model, and the parameters estimated from the calibration sites (table 1). c: the distance weighted vegetation proportions within 1800 m of the centre of the test sites, calculated from the a.d. 1800 land cover maps. cerealia calluna poaceae trees 0% 20% 40% 60% 80% 100% 0% 20% 40% 60% 80% 100% 0% 20% 40% 60% 80% 100% avn sø st or e g rib sø g ud m e sø g un ds øm ag le s ø d all er up s ø st or e ø ks sø sk ån sø n av ns ø kra gs ø a: pollen proportions c: vegetation from maps b: reconstructed vegetation portions of the samples. for example, the samples from all test sites contain more than 25% tree pollen, although the surroundings of many sites contain very little woodland, a fact which is reflected by the reconstructions. the amount of calluna tends to be higher in the reconstructions than estimated from the maps, which could indicate that the estimated αi for this species is too low, or that the a.d. 1800 heathlands were richer in calluna than was assumed based on present-day heathlands in southern sweden. however, the reconstructions clearly distinguish the sites where heathland was common (store økssø, skånsø, navnsø and kragsø; fig. 1). the reconstructed amounts of cerealia reflect very well the differences in the amount of arable land between the test sites. conclusions the use of historical maps has proved very useful for quantifying the relationship between pollen sedimentation and vegetation. it has provided an empirical validation of the prentice/sugita model of pollen dispersal for open cultural landscapes (nielsen 2004), made it possible to estimate the relevant source area of pollen, and provided insights into the factors affecting rsap (nielsen & sugita in press). finally, it has provided a set of erv model parameters and shown that erv-based reconstructions reflect the actual patterns of vegetation composition around the lakes (nielsen & odgaard 2005). the next step is to apply these findings to fossil pollen diagrams and reconstruct the cultural landscape of earlier periods, from which no maps are available. this research will continue at geus in the coming years, thanks to a grant from the carlsberg foundation. acknowledgements bent odgaard (university of aarhus, denmark) and peter rasmussen (geus) are thanked for providing pollen data from test sites, and shinya sugita (university of minnesota, usa) for access to unpublished computer programs. this is publication no. 10 in the project agrar 2000, funded by the danish research councils. it is also a contribution to the nordforsk network pollandcal, co-ordinated by m.-j. gaillard (university of kalmar, sweden). references broström, a. 2002: estimating source area of pollen and pollen productivity in the cultural landscapes of southern sweden – developing a palynological tool for quantifying past plant cover. lundqua thesis 46, 101–115. lund: university of lund. broström, a., sugita, s. & gaillard, m.-j. 2004: pollen productivity estimates for the reconstruction of past vegetation cover in the cultural landscape of southern sweden. the holocene 14, 368–381. bunting, m.j., gaillard, m.-j., sugita, s., middleton, r. & broström, a. 2004: vegetation structure and pollen source area. the holocene 14, 651–660. calcote, r. 1995: pollen source area and pollen productivity: evidence from forest hollows. journal of ecology 83, 591–602. davis, m.b. 2000: palynology after y2k – understanding the source area of pollen in sediments. annual review of earth and planetary sciences 28, 1–18. nielsen, a.b. 2003: pollen based quantitative estimation of land cover – relationships between pollen sedimentation in lakes and land cover as seen on historical maps in denmark a.d. 1800. danmarks og grønlands geologiske undersøgelse rapport 2003/57, 135 pp. nielsen, a.b. 2004: modelling pollen sedimentation in danish lakes around a.d. 1800 – an attempt to validate the pollscape model. journal of biogeography 31, 1693–1709. nielsen, a.b. & odgaard, b. 2004: the use of historical analogues for interpreting fossil pollen records. vegetation history and archaeobotany 13, 33–43. nielsen, a.b. & odgaard, b. 2005: reconstructing land cover from pollen assemblages from small lakes in denmark. review of palaeobotany and palynology 133, 1–21. nielsen, a.b. & sugita, s. in press: estimating relevant source area of pollen for small danish lakes around a.d. 1800. the holocene. odgaard, b. 1999: landbrugslandskabet gennem 2000 år. geologi – nyt fra geus 1, 4–6. odgaard, b.v. & rasmussen, p. 2000: origin and temporal development of macro-scale vegetation patterns in the cultural landscape of denmark. journal of ecology 88, 733–748. parsons, r.w. & prentice, i.c. 1981: statistical approaches to r-values and pollen-vegetation relationship. review of palaeobotany and palynology 32, 127–152. prentice, i.c. 1985: pollen representation, source area, and basin size: toward a unified theory of pollen analysis. quaternary research 23, 76–86. prentice, i.c. & parsons, r.w. 1983: maximum likelihood linear calibration of pollen spectra in terms of forest composition. biometrics 39, 1051–1057. sugita, s. 1993: a model of pollen source area for an entire lake surface. quaternary research 39, 239–244. sugita, s. 1994: pollen representation of vegetation in quaternary sediments: theory and method in patchy vegetation. journal of ecology 82, 881–897. sugita, s., macdonald, g.m. & larsen, c.p.s. 1997: reconstruction of fire disturbance and forest succession from fossil pollen in lake sediments: potential and limitations. in: clark, j.s. et al. (eds): sediment records of biomass burning and global change, 387–412. berlin: springerverlag. sugita, s., gaillard, m.-j. & broström, a. 1999: landscape openness and pollen records: a simulation approach. the holocene 9, 409–421. sutton, o.g. 1953: micrometeorology, 333 pp. new york: mcgraw-hill. 52 author’s address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: abn@geus.dk petroleum geochemistry of the deepened lopra-1/1a re-entry well, faroe islands 67© geus, 2006. geological survey of denmark and greenland bulletin 9, 67–77. available at: www.geus.dk/publications/bull petroleum geochemistry of the deepened lopra-1/1a re-entry well, faroe islands jørgen a. bojesen-koefoed and h. peter nytoft the lopra-1/1a re-entry well was drilled as a stratigraphic test with no immediate exploration objectives. hence, petroleum geochemical studies were of limited extent, and restricted to non-destructive analyses. the presence of natural petroleum hydrocarbons could not be confirmed with certainty, but hydrocarbons extracted from the hydrochloric acid solute of a calcite vug present in rswc #1 (3543 m), may represent indigenous petroleum since hydrocarbon-bearing fluid inclusions have been reported from the same sample. these hydrocarbons show some similarities to petroleum generated from the upper jurassic – lower cretaceous kimmeridge clay type source rocks present in surrounding areas. except for this sample, the results generally show the presence of a variety of contaminants of different origins such as ‘naturally greasy fingers’ (squalene and cholesterol), cosmetics such as chap stick or hand lotion (e.g. esters such as butyl-stearate, stearyl-palmitate, vitamin a), plasticisers (phthalates), diesel oil and ‘pipe dope’. keywords: oil traces, organic geochemistry, contamination, faroes, north atlantic, lopra ______________________________________________________________________________________________________________ geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: jbk@geus.dk as part of the preparatory activities prior to an expected future licensing round on faroese territory, a consortium of oil companies (see heinesen et al. 2006, this volume) undertook to drill the lopra-1/1a re-entry well as a stratigraphic test. the primary objectives of the well were to obtain lithological and stratigraphic information on the deepest parts of the faroese flood basalt sequence as well as its substratum, and to acquire information related to the assessment of the petroleum exploration prospectivity of the area. the original lopra-1 well was drilled in 1981 and reached a total depth (td) of 2178 m. the well penetrated a succession of flood basalts with minor tuffs and two dolerites (probably dykes). the drilled sequence was interpreted as part of the faroese lower basalt formation or series (waagstein et al. 1984; waagstein 1988). the lopra-1 well was re-entered in 1996 and deepened to a td of 3565 m. the new drilling showed that the subaerial flood basalt sequence was underlain by a subaqueous hyaloclastite-basalt succession, comprising the interval 2550–3565 m and consisting predominantly of lapilli-tuffs, tuff-breccias and beds of blocks of basalt. no clastic sedimentary rocks were recorded (boldreel 2006, this volume). the deepened lopra-1/1a re-entry well had no direct hydrocarbon exploration objectives and, as a technical consequence, other types of investigations took priority over petroleum geochemistry. this was particularly the case since no sedimentary rocks were penetrated and neither were any evident shows detected during drilling. accordingly, the analytical programme carried out with respect to petroleum geochemistry was somewhat limited in scope and character. previously, various indications had been recorded of the presence of hydrocarbons in the original lopra-1 well prior to deepening (jørgensen 1984; laier et al. 1997): 1. during drilling in 1981, a gas show was noted at 2008 m brkb and ‘wax’/bitumen was observed on zeolites. 2. in 1983, gas at a pressure of 19 bar was sampled at the wellhead. the gas flowed at approximately 1 m3 per day and consisted of methane (72%), n2 (27%), and traces of higher hydrocarbons (jacobsen & laier 1984). geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1967 68 in addition, water with an oil-film was produced by the well, but the oil film was not analysed in detail. based on its isotopic composition, it was concluded that the gas was principally of thermogenic origin, although the gas was unusually ‘dry’ (laier et al. 1997). 3. in 1992, oil-film on water from the well was sampled, and minor amounts of gas escaped from the wellhead. based on geochemical analysis, it was suggested that the oil originated from a mature siliciclastic source rock, which was presumed to be present below the basaltic cover (laier et al. 1997). 4. during a vsp-survey in 1994, a logging tool was lowered into the well and a thick, oil-smelling mud-slurry was observed sticking to the tool and wire when returned to the surface. minor amounts of gas were noted. analysis of the slurry revealed the presence of petroleum components similar to those detected in water samples from the well. moreover, wax-coatings on zeolite minerals have been recorded in several outcrops on the faroe islands (jørgensen 1984; laier et al. 1997). in order to check if these indications of the presence of petroleum in the lopra-1 well prior to deepening could be further substantiated, a minor petroleum geochemical analytical programme was carried out on samples from the deepened well, the results of which are reported below. samples and methods since the lopra-1/1a re-entry well was drilled as a stratigraphic test, samples were principally reserved for investigations directly related to the main objectives of the programme, and petroleum geochemical studies were generally restricted to non-destructive analyses. hence, samples could not be crushed and analyses had to rely on organic matter extracted from the surface and the directly accessible pore-spaces of the recovered rock samples. this procedure is necessarily sub-optimal, since, as will be demonstrated, the risk of contamination is overwhelming. however, for pre-determined reasons other types of investigations had been assigned greater priority. all analyses were carried out on rotary sidewall cores (rswc), 58 of which were collected over the section penetrated in the deepened well below the td of the original lopra-1 well, i.e. deeper than 2178 m brkb. all samples consisted of basaltic volcanics, i.e. lavas, tuffs and hyaloclastic breccias. upon receipt, all rswc samples, which came in screwcap glass containers, were taken to a darkroom and checked by organolfactoric means for petroleum odour, and for visible fluorescence by means of a hand-held uv-lamp. based on the presence of a distinct petroleum odour and weak to clear fluorescence, seven samples were selected for further study (table 1). table 1. sample identification, solvent extraction and gas chromatographic key data sample rswc #51 rswc #45 rswc #36 rswc #33 rswc #24 rswc #13 rswc #1 rswc #1 rswc #1 drilling mud drilling mud pipe dope a in wt% of total extract. b in wt% of maltene fraction. c pristane/phytane ratio, from gas chromatography. d carbon preference index, calculated over the interval nc22–32. depth (m b. kb) 2361 2450 2570 2630 3076 3438 3543 3543 3543 2900 3200 vesicular basalt basalt welded tuff welded tuff basalt lapillic tuff tuff w. calcite vug tuff w. calcite vug tuff w. calcite vug entire rswc immersed in dcm entire rswc immersed in dcm entire rswc immersed in dcm entire rswc immersed in dcm entire rswc immersed in dcm entire rswc immersed in dcm entire rswc immersed in dcm calcite vug dissolved in hcl, organic extract recovered by shaking with dcm tuff chip, counterpart of calcite vug, crushed and extracted approximately 10 g of drilling mud before addition of diesel, rinsed by dcm approximately 10 g of drilling mud after addition of diesel, rinsed by dcm anti-seize compound for drilling rods sample supplied by danop extract recovery (mg) 1.0 0.8 0.6 0.8 0.7 1.1 n.d. 1.1 6.3 5.7 23.9 n.a. asphaltenesa n.a. n.a. n.a. n.a. n.a. n.a. n.a. 27.3 76.2 7.0 2.5 2.1 saturatesb n.a. n.a n.a. n.a. n.a. n.a. n.a. n.a. 22.2 67.3 68.0 57.8 aromaticsb n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. 11.1 19.2 22.2 24.5 nsob n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. 66.7 13.5 9.9 17.7 pr/phc 1.61 1.50 1.52 1.41 1.36 1.34 1.38 1.02 1.20 1.15 1.66 n.a. cpid 1.11 1.17 n.a. n.a. n.a. n.a. n.a. 1.24 1.33 1.17 1.05 n.a. commentlithology rswc: rotary sidewall core. dcm: dichloromethane. n.a.: not analysed. n.d.: not detected. kb: kelly bushing. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1968 69 in order to preserve the samples intact for other investigations, extracts were recovered from the surface and the immediately accessible pore spaces of the samples by immersing entire rswcs into dichloromethane (dcm) for approximately 15 minutes in a glass beaker placed in an ultrasonification device. rock fragments were separated by centrifugation, dcm was removed by evaporation, and the recovery was determined by weighing. due to very low recoveries, no further preparative procedures were applied. total extracts were re-dissolved in isooctane and analysed by gas chromatography (gc) and by coupled gas chromatography – mass spectrometry (gcms). rswc #1 (3543 m) was treated as described above, but a small chip containing a calcite-filled vug was removed from the sample. the vug was dissolved in dilute hydrochloric acid (hcl, 2n), and an organic extract was recovered by several stages of shaking the solute with dichloromethane (dcm) in a separatory funnel (table 1). the basaltic counterpart of the vug was coarsely crushed and extracted for four hours in a soxtec apparatus (1h reflux in dcm, followed by 3h rinsing). asphaltenes were precipitated from both extracts by addition of 40-fold excess of n-pentane, and the maltenes were separated into saturated, aromatic and polar compounds by medium pressure liquid chromatography (mplc), using a method modified from radke et al. (1980). saturate fractions were analysed by gc and gc-ms. in order to remedy problems with a stuck pipe, diesel was added to the well at a depth of approximately 3100 m. neither samples of the diesel, nor of the drilling mud, were preserved for analysis. instead, drilling mud from bagged drill cutting samples collected before and after addition of diesel were analysed in order to check for possible contamination. approximately 10 g of sample (drilling mud plus cuttings) were ultrasonically extracted with dcm (app. 100 ml). extracts were recovered by decantation and centrifugation, and the dcm removed by evaporation. asphaltenes were precipitated by addition of 40fold excess of n-pentane, with the maltenes being separated into saturated, aromatic and polar components as described above. the saturate fractions were analysed by gc and gc-ms. an additional possible source of contamination was pipe dope, an anti-seize compound used when joining drilling rods. a sample of the pipe-dope used during drilling of the lopra-1/1a re-entry well was supplied by danop and analysed using standard procedures for oil analysis. asphaltenes were precipitated by addition of 40-fold excess of n-pentane, and the maltenes were separated into saturated, aromatic and polar components as described above. again the saturate fractions were analysed by gc and gc-ms. gas chromatographic analyses were carried out by means of a hewlett-packard 5890 series ii plus gas chromatograph, using splitless injection, a 25 hp-1 wcot column and a flame ionisation detector (fid). biological marker analyses were carried out by means of a hewlett-packard 5980a series ii gas chromatograph interfaced to a hewlett-packard 5971 mass selective detector (msd) using splitless injection and a 25 m hp-5 wcot column. rswc #51, 2361 m rswc #45, 2450 m rswc #36, 2470 m rswc #13, 3438 m rswc #1, 3543 m rswc #33, 2630 m rswc #24, 3076 m, basalt diesel non-diesel fig. 1. gas chromatograms obtained from analysis of rswcs rinsed in dcm. approximate shapes and positions of ‘diesel’ and ‘nondiesel’ envelopes are shown in sample rswc #51. examples of ‘prominent unknowns’ are indicated by asterisks in sample rswc #36, see fig. 3. note that all samples show some degree of diesel contamination, although this adulterant was not added until a depth of approximately 3100 m. this observation suggests that diesel contamination is pervasive throughout the entire uncased section of the well. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1969 70 results rswcs rinsed in dcm total extract recoveries were generally close to 1 mg, and no attempt was made to fractionate the residues. gas chromatograms of the bulk extracts (fig. 1) show that most samples contain a more or less well-developed series of n-alkanes. the distributions are generally light-end skewed and unimodal, but two samples show evidence of bimodal distributions with higher proportions of waxy (22+) n-alkanes. pristane/phytane ratios range from 1.34 to 1.61, and no oddor even-number predominance is noted among the n-alkanes. in addition to n-alkanes, a number of prominent peaks of unknown identity are noted in all but one of the chromatograms, see below. key parameters are listed in table 1. biological marker maturity indications are consistent, with homohopane and bishomohopane 22s/(22s + 22r) epimerisation ratios at equilibrium (i.e. close to 0.60), and c29 sterane 20s/(20s + 20r) epimerisation ratios slightly below equilibrium (i.e. slightly less than 0.52), indicating early to mid oil-window maturity (table 2). sterane and triterpane biological marker distributions display only minor variations among the samples. representative ion fragmentograms m/z 191, m/z 217, and m/z 218 are shown in fig. 2, with key biological marker parameter ratios tabulated in tables 3, 4. triterpane distributions generally comprise fairly high proportions of tricyclic triterpanes. among the pentacyclics the presence of 30-norhopanes plus 25norhopanes is noted together with a peak eluting fractionally earlier than hopane. this peak is routinely assigned to 18α(h)-oleanane, a well-known marker of angiosperm higher plant inputs, and hence of source ages younger than the mid cretaceous. however, scan-mode mass spectrometric analysis could not confirm the identity of this compound. rather the peak represents several co-eluting compounds, probably comprising oleanane, lupane, and one or more unknowns. limited amounts of sample precluded further investigation of this problem of identification. norhopane to hopane ratios are close to unity and extended hopanes are relatively abundant. moreover, gammacerane and notable proportions of hexahydrobenzohopane are present. sterane distributions are very similar for all samples, comprising a clear predominance of c27 steranes over the c28 plus c29 steranes, but with the co-occurrence of c30 as well as of c26 steranes. prominence of ββ-sterane epimers is noted in both the m/z 217 and m/z 218 ion fragmentograms. a number of prominent unknowns are noted in the pr is ta ne ph yt an e n15 n20 n25 squalene m/z 218 sq 29 28 27 26 25 m/z 191 sq 1 2 3 4 5 6 7 8 9 10 11 12 15 16 19 20 21 22 23 13 14 1718 m/z 217 sq 25 27 29 28 s r 26 24 * * * * fig. 2. gas chromatogram and representative ion fragmentograms m/z 191, m/z 217 and m/z 218 (sample rswc #45, 2450 m). filled black peak labelled ‘sq’ is the contaminant squalene – a compound found on the skin of humans. compound identification is shown in table 5. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1970 71 gas chromatograms as described above. scan-mode mass spectrometric analyses of the extract obtained from rswc #1, in which these compounds are particularly abundant, identifies: squalene and cholesterol; 2-ethylhexyl-phthalate; vitamin a and various esters such as butyl-stearate, stearyl-palmitate and a series of similar compounds (fig. 3). calcite vug dissolved in hcl and basalt coarsely crushed total extract recovery of the acid-digested calcite vug was only 1.1 mg, whereas solvent extraction of the tuff hostrock yielded 6.3 mg. the comparatively high extraction yield of the latter sample is probably due to the more efficient extraction procedure, i.e. crushing and soxtec extraction as opposed to ultrasonical extraction of entire samples. gas chromatograms of saturate fractions of the two sub-samples (fig. 4) are different. the extract recovered from the acid-digested calcite vug shows a strongly unimodal n-alkane distribution, centred around c20, with notable light-end depletion and a rather poor signal-tonoise ratio. the basalt extract shows a well-developed unimodal distribution of n-alkanes, centred around c17, and a clear odd-predominance in the c25–32 range. pristane/phytane ratios are 1.02 and 1.20, respectively. biological marker distributions are similar in the two samples, although the signal-to-noise ratio observed in the acid-digested sample is rather poor (fig. 4). these distributions are significantly different from the picture provided by other samples, including both sidewall cores (see above), drilling mud and pipe dope (see below). triterpane distributions comprise low proportions of tricyclic triterpanes. amongst the pentacyclics the presence of 28,30-bisnorhopane is noted, while 30-norhopanes, 25norhopanes, and ‘oleanane’ are absent, or cannot be identified with any degree of certainty. h29/h30 ratios are close to 0.3 and extended hopanes are relatively scarce. regular sterane distributions comprise a clear predominance of c27 steranes over the c28 plus c29 homologues with the presence of c30 as well as c26 steranes. drilling mud total extract recovery from the drilling mud samples differs widely before and after the addition of diesel to the drilling mud at a depth of approximately 3100 m brkb (table 1). this difference recurs in the gas chromatograms of saturate fractions of the two samples (fig. 5). the sample collected before addition of diesel yields a rather irregular light-end skewed n-alkane distribution with high proportions of ‘unresolved complex mixture’ (ucm). the sample collected after addition of diesel yields a well-defined, nearly symmetrical, unimodal n-alkane distribution, centred around c16. pristane/phytane ratios are 1.15 and 1.66, respectively. except for a minor enhancement of c27 diasteranes, and a slightly more pronounced enhancement of low molecular weight tricyclic triterpanes in the diesel-containing sample, biological marker distributions, however, are similar in the two samples and indistinguishable from time (min.) 10 20 30 40 50 60 abundance 0 2.4 x 107 2.0 x 107 pr is ta ne ph yt an e ? ? ? ? ? ? st ea ry l-p al m ita te c ho le st er ol squalene bu ty l-s te ar at e v ita m in -a ‘diesel-envelope’ common natural contaminants from ‘greasy fingers’ various hand lotion and chap stick components plasticiser for e.g. pvc plastic 1.6 x 107 1.2 x 107 8 x 106 4 x 106 2-ethylhexyl-phthalate + + + + + + + + +++x x * * * fig. 3. full scan total ion fragmentogram, rswc #1. the approximate shape and position of a ‘diesel envelope’ are shown by shading. interpretation of the origins of various contaminants is shown. s29 s/(s+r)a 0.49 0.48 0.46 0.43 0.48 0.49 0.47 0.44 0.51 0.45 0.50 0.42 sample rswc #51 rswc #45 rswc #36 rswc #33 rswc #24 rswc #13 rswc #1 rswc #1 rswc #1 drilling mude drilling mudf pipe dope a c29 regular sterane αα20s/(αα20s+αα20r) epimer ratio. b c29 regular sterane ββ/(ββ+αα) epimer ratio. c homohopane 22s/(22s+22r) epimer ratio. d bishomohopane 22s/(22s+22r) epimer ratio. e before addition of diesel. f after addition of diesel. rswc: rotary sidewall core. s29 βα/(ββ+αα)b 0.62 0.60 0.60 0.63 0.61 0.60 0.60 0.58 0.54 0.61 0.64 0.42 table 2. biological marker maturity data 0.61 0.59 0.60 0.59 0.56 0.57 0.59 0.58 0.58 0.59 0.59 0.65 h31 s/(s+r)c 0.63 0.61 0.60 0.63 0.59 0.60 0.59 0.59 0.61 0.62 0.59 0.57 h32 s/(s+r)d geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1971 72 the biological marker distributions yielded by rswcs rinsed in dcm (fig. 5). pipe dope the gas chromatogram of the saturate fraction does not allow identification of any components, but simply shows a large hump of ‘unresolved complex mixture’ (fig. 6). biological marker maturity parameters indicate early to mid-oil window maturity (table 2). triterpane distributions generally comprise high proportions of tricyclic triterpanes, and among the pentacyclics the presence of 30-norhopanes plus abundant 25-norhopanes is noted together with a peak eluting fractionally earlier than hopane, probably representing several co-eluting compounds. these possibly comprise oleanane, lupane and one or more unknowns (fig. 6). the h29/h30 ratio is 1.12 with extended hopanes being abundant. gammacerane and notable proportions of hexahydrobenzohopane are present. the distribution of regular steranes shows a clear predominance of c27 steranes over the c28 and c29 homologues plus the presence of c30 as well as c26 steranes. prominence of ββ-sterane epimers is noted in both the m/z 217 and m/z 218 ion fragmentograms. discussion the amounts of extract recovered from the rswc samples are generally very low, which naturally limits the possibilities for detailed studies and implementation of extensive sample preparation techniques in order to optimise the quality of analytical data obtained. furthermore, since total recoveries are low, even minor random contamination, which normally would be insignificant and ignored, may cause notable problems. hence, considerable uncertainty is attached to the conclusions made on the basis of the analyses reported here. analyses of mud samples collected before and after addition of diesel show a profound influence of diesel on the content and distribution of n-alkanes, whereas the biological marker characteristics, except for minor enhancement of c27 diasteranes and low molecular weight tricyclic triterpanes, are largely unaffected. hence for most practical purposes, this particular diesel distillate fraction will not severely influence any of the biomarker ratios. the biomarker distributions observed in the two mud samples are largely identical to the distributions yielded by dcm-rinsed rswc samples. based on gas chromatography data, most dcm-rinsed rswc samples are seen to contain diesel, but some extracts also contain longer chain-length n-alkanes and biological markers which are unlikely to originate from diesel contamination. in addition to diesel, contamination from various other sources is present in most samples: sample rswc #51 rswc #45 rswc #36 rswc #33 rswc #24 rswc #13 rswc #1 rswc #1 rswc #1 drilling mudf drilling mudg pipe dope a c23 tricyclic terpane to hopane ratio. b ts: 18α(h)-trisnorneohopane, tm: trisnorhopane. c h28: 28,30-bisnorhopane, h29: norhopane. d h29: norhopane, h30: hopane. e ol: 18α(h)-oleanane, h30: hopane. f before addition of diesel. g after addition of diesel. n.a.: not analysed. rswc: rotary sidewall core. table 3. key triterpane biological marker parameter ratio 0.00 0.14 0.18 0.14 0.52 0.80 0.36 0.00 0.00 0.03 0.04 0.16 ol/h30e 0.96 1.06 0.96 0.93 1.01 0.96 0.94 0.33 0.28 0.99 1.02 0.12 h29/h30d 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.35 0.37 0.00 0.00 0.00 h28/h29c 0.49 0.52 0.52 0.54 0.57 n.a. n.a. 0.43 0.44 0.51 0.51 0.41 ts/(ts+tm)b 0.93 1.00 1.00 1.14 1.00 1.06 0.64 0.13 0.12 0.50 1.28 0.44 t23/h30a a (sum c27 diateranes)/(sum c27 regular steranes), m/z 217. b relative distribution of c27–29 regular steranes, based on αα20r epimers in m/z 217. c c27/c29 regular sterane ratio, based on αα20r epimers in m/z 217. d c30 regular steranes. e before addition of diesel. f after addition of diesel. rswc: rotary sidewall core. sample rswc #51 rswc #45 rswc #36 rswc #33 rswc #24 rswc #13 rswc #1 rswc #1 rswc #1 drilling mude drilling mudf pipe dope d27/s27a 1.33 1.27 1.39 1.53 1.20 1.22 0.95 1.04 1.06 0.81 1.19 0.47 table 4. key sterane biological marker parameter ratios 19.0 21.3 23.0 20.0 19.56 21.7 22.4 21.3 21.4 19.5 24.6 23.8 s28 (%)b 30.2 32.0 31.0 30.0 29.4 30.4 30.3 31.9 30.4 33.5 29.5 32.8 s29 (%)b 1.7 1.5 1.5 1.7 1.7 1.6 1.6 1.5 1.6 1.4 1.6 1.3 s27/s29c present present present present present present present present present present present present s30d 50.8 46.7 46.0 50.0 51.0 47.8 47.4 46.8 48.2 47.0 45.9 43.3 s27 (%)b geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1972 73 pr is ta ne ph yt an e n15 n20 n25 pr is ta ne ph yt an e n15 n20 n25 m/z 191 28 ,3 0bi sn or ho pa ne 3 4 9 10 11 15 16 17 19 20 21 22 13 m/z 191 28 ,3 0bi sn or ho pa ne 3 4 9 10 11 15 16 19 20 21 22 13 m/z 218 28 2726 25 m/z 217 24 25 27 28 s r 26 * * * * m/z 218 28 29 2726 25 m/z 217 24 25 27 28 s r 26 * * * * fig. 4. gas chromatograms and ion fragmentograms m/z 191, m/z 217 and m/z 218, rswc #1, calcite vug dissolved in hcl (left), and its coarsely crushed tuff host rock (right). note mutual similarity of biological marker distributions and differences when compared to distributions shown in figs 2, 5, 6. compound identification is shown in table 5. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1973 74 m/z 218 m/z 218 29 2827 26 25 29 28 27 26 25 m/z 191 1 2 3 4 5 6 7 8 9 10 11 12 15 16 19 20 21 22 23 13 14 17 18 m/z 217 25 27 28 29 s r 26 24 * * * * m/z 191 1 2 3 4 5 6 7 8 9 10 11 12 15 16 19 20 21 22 23 13 14 17 18 m/z 217 25 27 28 29 s r 26 24 * * * * ucmpr is ta ne ph yt an e n15 n20 n25 pr is ta ne ph yt an e n15 n20 n25 fig. 5. gas chromatograms and ion fragmentograms m/z 191, m/z 217 and m/z 218, extract of drilling mud before addition of diesel (left), and extract of drilling mud after addition of diesel (right). ucm, unresolved complex mixture. note profound influence of diesel on the n-alkane distribution, and the lack of, or limited effect on, biological marker distributions. compound identification is shown in table 5. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1974 75 squalene and cholesterol are commonly occurring natural compounds present on the skin of humans, for instance on the hands (‘naturally greasy fingers’); 2-ethylhexyl-phthalate is a widely used plasticiser for various polymers; vitamin a, esters such as butyl-stearate, stearyl-palmitate and similar compounds detected in the sample are commonly used in cosmetics, including hand lotion and chap-stick. the absence of both acyclic isoprenoids and n-alkanes from pipe dope plus the presence of abundant 25-norhopanes suggest that this product is based on a heavily biodegraded oil (corresponding to level 6 of peters & moldowan (1993)). based on the presence of 30-norhopanes, which is also manifest in norhopane to hopane ratio close to unity plus the prominence of ββ-sterane epimers and comparatively high proportions of homohopanes (in particular tetrakishomohopane to pentakishomohopane ratio close to unity), this oil was presumably generated from a marine marly source rock, deposited in a highly anoxic environment. in addition, the presence of angiosperm higher land plant markers such as 18α(h)-oleanane suggest a source age younger than the mid-cretaceous. normal and acyclic isoprenoid alkanes are absent, but the biological marker distribution shows clear similarities to the distributions yielded by dcm-rinsed rswcs discussed above, such that the presence of contamination from pipe dope, in addition to adulteration from other sources, seems obvious. however, differences are observed: pipe dope contains very high proportions of 25-norhopanes, whereas the proportion in the rswcs are but minor; the relative abundance of ts and tm is reversed in pipe dope compared to dcm-rinsed rswc samples. similarly, the proportion of c27 diasteranes relative to c27 regular steranes is much lower in pipe dope extracts than in the dcmrinsed rswc samples. the latter feature may, however, be wholly or partly caused by the addition of diesel, which was shown above to result in minor enhancement of low boiling-range tricyclics and diasteranes relative to noncontaminated samples. in summary, dcm-rinsed samples are contaminated by a variety of compounds originating from several sources, including diesel, pipe dope, plasticisers, naturally greasy fingers and cosmetics, possibly hand lotion and/or chap stick. however, the samples also contain petroleum components that do not seem to originate from these sources of contamination. an origin from other sources of contamination or from indigenous crude oil is conceivable. laier et al. (1997) show the presence of traces of heavy petroleum hydrocarbons and wax in samples from the original lopra-1 well, prior to deepening. the ‘unexplainable’ petroleum components found in dcm-rinsed samples from the deepened well may represent similar occurrences. level n15 a b n20 n25 sq n30 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 24 25 26 27 28 29 c15 normal alkane pristane phytane c20 normal alkane c25 normal alkane squalene c30 normal alkane c28 tricyclic terpanes (2 isomers) c29 tricyclic terpanes (2 isomers) ts = trisnorneohopane tm = trisnorhopane c30 tricyclic terpanes (2 isomers) c28 25,30 bisnorhopane, coeluting with "5" c31 tricyclic terpanes (2 isomers) c29 25-norhopane, partially coeluting with "7" norhopane = 30-norhopane 29ts = norneohopane normoretane mixture of oleanane, lupane and unknown hopane c30 30-norhopane moretane homohopane, 22s and 22r isomers gammacerane c31 hexahydrobenzohopane homomoretane bishomohopane, 22s and 22r isomers trishomohopane, 22s and 22r isomers tetrakishomohopane, 22s and 22r isomers pentakishomohopane, 22s and 22r isomers c27 diasteranes, 4 isomers labelled with asterisks c26 regular steranes, ββ isomers only c27 regular steranes, αα and ββ isomers c28 regular steranes, ββ isomers only c29 regular steranes, ααs, ααr and ββ isomers c30 regular steranes, ββ isomers only table 5. compound identification key compound biological marker, in particular triterpane, distributions in extracts recovered from dcm-rinsing of rswc samples from the deepened lopra-1 well show a number of striking similarities to distributions yielded by certain samples collected in the original lopra-1 well prior to re-entry. these samples include petroleum extracted from water samples in 1992 (laier et al. 1997), and a sample collected from slurry sticking to a vsp-tool, which was lowered into the hole in 1994. the biological marker characteristics include features such as the presence of 30-norhopanes, 25-norhopanes, gammacerane, hexahydrobenzohopane, h29/h30 ratios close to or greater than 1 and a high relative abundance of extended hopanes plus prominence of the ββ-sterane epimers. minor differences between pipe dope and the ‘slurry’ and the oil film are also observed, principally with respect to the presence of n-alkanes and the proportions of c28 regular steranes relative to the progeus bulletin no 9 7 juli.pmd 07-07-2006, 14:1975 76 gest the presence of pipe dope contamination in the original borehole as well, the differences being accounted for by presumed differences between the pipe dopes used in the original and the deepened lopra wells. furthermore, the presence of a mid-cretaceous or younger marly, anoxic marine source rock in the area, as implied by the geochemical data, seems geologically problematical. carbonate or marly source rocks typically occur in lower latitude regions, i.e. areas within and close to the arid tropical belts (tissot & welte 1984). it is estimated that during the cretaceous, the faroe islands area was situated at 40–45°n, and northward movement has prevailed since then (habicht 1979; scotese et al. 1988). the biological marker characteristics of the extracts recovered from acid-digestion of a calcite-filled vug, and from its crushed tuff host-rock are totally dissimilar to the characteristics shown by all other samples from the well, including pipe dope. principally, the distribution of triterpanes serves to distinguish these two samples from the remainder of the samples analysed. hence, the following characteristic are noteworthy: lower proportions of tricyclic triterpanes, low norhopane to hopane ratio and the presence of 28,30-bisnorhopane. the signal-to-noise ratio is comparatively poor, but the overall biological marker characteristics show some similarities to oils generated from upper jurassic kimmeridge clay formation sediments and their equivalents. this source system is known to be present in west of shetlands basins (e.g. scotchman et al. 1998) as well as in the north sea basins, and can be prognosed for faroese waters. glassley (2006, this volume) estimates that the maximum temperature reached at td of the lopra-1/1a reentry well was 200°c. provided that this estimate is correct, the temperature may be too high to allow preservation of higher molecular weight liquid hydrocarbons if maintained over prolonged periods of time. the maximum temperature actually recorded in the well was 98°c, and assuming that the hydrocarbons found in rswc #1 represent a thermogenic natural petroleum, this may have entered the tuff at a temperature lower than the maximum estimated by glassley (2006, this volume), probably in connection with the formation of the calcite vug. petroleum-bearing fluid inclusions have been observed in the same sample (konnerup-madsen 2006, this volume). based on fluorescence colours (orange-yellow to green), an api gravity of 20–35 is estimated. no homogenisation temperature data have been recorded for the petroleumbearing inclusions, but data from non-hydrocarbon bearing inclusions fall in the range 101–186°c, with the higher temperatures probably being caused by partial decrepitation. hence, entrapment temperatures were probably con1 2 3 4 5 6 7 8 9 10 11 12 15 16 19 20 21 22 23 25 27 29 29 28 27 26 25 28 s r 26 13 14 17 18 ucm 24 * * * * m/z 191 m/z 217 m/z 218 fig. 6. gas chromatogram and ion fragmentograms m/z 191, m/z 217 and m/z 218, pipe dope. ucm, unresolved complex mixture. compound identification is shown in table 5. portions of c27 and c29 regular steranes. however, the striking similarities and the presence of somewhat unusual components in pipe dope, as well as in the samples collected from the lopra-1 well, prior to deepening, may suggeus bulletin no 9 7 juli.pmd 07-07-2006, 14:1976 77 siderably lower than the maximum temperature as estimated by glassley (2006, this volume), and it is conceivable that the signal obtained represents indigenous crude oil. if so, this observation is very encouraging for future exploration in faroese territorial areas. conclusions all samples are to variable degrees contaminated by a number of adulterants of different origins. the adulterants include: 1. n-alkanes and other petroleum components originating from commercial diesel fuel; 2. pipe dope, an anti-seize compound used, for instance, when joining drilling rods; 3. squalene and cholesterol originating from naturally greasy fingers; 4. vitamin a and various esters used in cosmetics, including chap stick and hand lotion; 5. 2-ethylhexyl-phthalate, a widely used plasticiser for polymers/plastics. a number of striking similarities in biological marker distributions between pipe dope and samples collected in the well prior to re-entry and deepening may suggest the presence of pipe dope contamination in the original well too, in addition to the presence of traces of petroleum hydrocarbons as shown by laier et al. (1997). organic extracts recovered from dissolution of a calcite vug in rswc #1 (3543 m) and from its tuff host rock yield biological marker distributions different from all other samples collected in the lopra-1 well. the distribution hints at generation from a marine anoxic shale source rock similar to the kimmeridge clay formation and equivalents known from surrounding areas. it is conceivable that this organic extract represents an indigenous thermogenic petroleum, in particular since oil-bearing fluid inclusions have been observed in the same sample. acknowledgements danop kindly supplied a sample of the pipe dope used during drilling of the well. ditte kiel-dühring assisted in the preparation and analysis of the samples. troels laier and the reviewers dr. r. burwood and dr. g. van graas provided helpful comments and suggestions, which significantly improved the manuscript. references boldreel, l.o. 2006: wire-line log-based stratigraphy of flood basalts from the lopra-1/1a well, faroe islands. geological survey of denmark and greenland bulletin 9, 7–22 (this volume). glassley, w.e. 2006: mineralogical and thermodynamic constraints on palaeogene palaeotemperature conditions during low-grade metamorphism of basaltic lavas recovered from the lopra-1/1a deep hole, faroe islands. geological survey of denmark and greenland bulletin 9, 109–118 (this volume). habicht, j.k.a. 1979: paleoclimate, paleomagnetism, and continental drift. aapg studies in geology 9, 31 pp. + maps. heinesen, m.v., larsen, a.r. & sørensen, k. 2006: introduction. geological survey of denmark and greenland bulletin 9, 5–6 (this volume). jacobsen, o.s. & laier, t. 1984: analysis of gas and water samples from the vestmanna-1 and lopra-1 wells, faeroe islands. in: berthelsen, o., noe-nygaard, a. & rasmussen, j. (eds): the deep drilling project 1980–91 in the faeroe islands. annales societatis scientiarum færoensis supplementum ix, 149–155. tórshavn: føroya frodskaparfelag. jørgensen, o. 1984: zeolite zones in the basaltic lavas of the faeroe islands. in: berthelsen, o., noe-nygaard, a. & rasmussen, j. (eds): the deep drilling project 1980–91 in the faeroe islands. annales societatis scientiarum færoensis supplementum ix, 71– 91. tórshavn: føroya frodskaparfelag. konnerup-madsen, j. 2006: a reconnaissance study of fluid inclusions in fracture filling quartz and calcite from the lopra-1/1a well, faroe islands. geological survey of denmark and greenland bulletin 9, 119–122 (this volume). laier, t., nytoft, h.p., jørgensen, o. & isaksen, g.h. 1997: hydrocarbon traces in the tertiary basalts of the faeroe islands. marine and petroleum geology 14, 257–266. peters, k.e. & moldowan, j.m. 1993: the biomarker guide, 363 pp. new jersey: prentice hall. radke, m., willsch, h. & welte, d.h. 1980: preparative hydrocarbon group type determination by automated medium pressure liquid chromatography. analytical chemistry 52, 406–411. scotchman, i., griffith, c.e., holmes, a.j. & jones, d.m. 1998: the jurassic petroleum system north and west of britain: a geochemical oil–source correlation study. organic geochemistry 29, 671–700. scotese, c.r., gahagan, l.m. & larson, r.l. 1988: plate tectonic reconstruction of the cretaceous and cenozoic ocean basins. tectonophysics 155, 27–48. tissot, b. & welte, d.h. 1984: petroleum formation and occurrence, 2nd edition, 699 pp. berlin: springer verlag. waagstein, r. 1988: structure, composition and age of the faeroe basalt plateau. in: morton, a.c. & parson, l.m. (eds): early tertiary volcanism and the opening of the ne atlantic. geological society special publication (london) 39, 225–238. waagstein, r., hald, n., jørgensen, o., nielsen, p.h., noe-nygaard, a., rasmussen, j. & schönharting, g. 1984: deep drilling on the faeroe islands. bulletin of the geological society of denmark 32, 133–138. manuscipt received 16 june 2000; revision accepted 9 november 2000. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1977 geological survey of denmark and greenland bulletin 41, 2018, 91-94 91 fish otoliths, also called ear stones or statoliths, are calcified structures functioning as movement and equilibrium indicators in the inner ear of fish (fig. 1). from hatching to death these structures grow incrementally, with new material accreted daily (pannella 1971) in successive layers of protein (1–8%, degens et al. 1969) and calcium carbonate. the accretion rate of otoliths varies with fish growth, and in temperate species it is usually lowest during the winter season (hüssy et al. 2010). this results in concentric growth resembling the ringed structure in trees (fig. 1d), enabling the use of dendrochronological techniques to approximate the age and growth history of fish. during growth, certain elements are incorporated into the otolith structure, some associated with proteins and some with the calcium carbonate component (thomas et al. 2017), supplying a valuable record of different aspects in fish life history and serving as a potential environmental record. previous studies show that trace element and isotopic compositions of otoliths can be used as a proxy for reconstructing water chemistry, temperature and salinity (patterson et al. 1993; thorrold & shuttleworth 2000). other studies demonstrate that elemental histories can be used to investigate fish spawning and migration patterns (e.g. sturrock et al. 2012), and more recent studies use elements such as zn, cu and mg as indicators of seasonality (hüssy et al. 2016; limburg et al. 2018). combining this knowledge of elemental variation with the micro-beam capabilities of laser ablation inductively coupled plasma mass spectrometry (laicpms) turns otolith microchemistry into a powerful tool for studying important parameters fundamental for establishing modern, sustainable fisheries management policies (e.g. stock identification, migration, pollution indicators, spawning habitats, duration of larval and juvenile stages, and magnitude and timing of spawning). we present an analytical method developed by the geological survey of denmark and greenland (geus) in collaboration with the national institute of aquatic resources, technical university of denmark (dtu aqua), for element abundance analysis in otoliths. analyses of otoliths from baltic cod (gadus morhua; fig. 1) are used as an example for its application. analytical approach the microchemical analysis of otoliths focuses on mg, p, ca, mn, cu, zn, sr and ba, as these elements are typically incorporated into otoliths, and are either subject to environmental control (e.g. sr and ba correlate with water salinity) or physiological control (e.g. zn, cu and mg are useful to the interpretation of otolith growth history; hüssy et al. 2016 and references therein; limburg et al. 2018). the la-icpms facility at geus employs a nwr213 laser system coupled to an element 2 double-focusing, single-collector magnetic sector field icpms. operating conditions, data acquisition and processing parameters are listed in table 1. la-icpms is often the preferred technique for rapid, in-situ analyses of trace elements and isotopes obtained from natural samples, simon hansen serre, kristian ege nielsen, peter fink-jensen, tonny bernt thomsen and karin hüssy analysis of cod otolith microchemistry by continuous line transects using la-icp-ms 5 mm c d a 2 mm b fig. 1. a: cod specimen caught in the baltic sea. b: removal of otoliths, the cut is situated just above the eyes. c: otolith, with dotted line showing the direction of a cross-section. d: photograph (reflected light) of a polished cross-section of an otolith. the red line shows the position of the laser scan. © 2018 geus. geological survey of denmark and greenland bulletin 41, 91–94. open access: www.geus.dk/bulletin http://www.geus.dk/bulletin 9292 as it requires little sample preparation and can produce high sample throughput, extracting elemental and isotopic information at a micrometre scale. most conventional la-icpms analysis is performed by spot analyses, following a bracketing analysis protocol using well-characterised standard materials. this is a powerful method when studying specific areas in solid materials. however, for compositional variations along millimetreto centimetre-scale transects, the spot approach becomes very time-consuming. for example, a 5 mm long transect requires about one hundred spots 40 µm in diameter, taking 3–4 hours to complete. spot analysis also introduces difficulties like downhole-element fractionation when the laser drills into the material. instead, research of fish otolith microchemistry favours faster sampling approaches such as line scans across the sample to acquire continuous timeresolved compositional profiles (a 5 mm long line takes c. 17 min, using a fixed scan speed of 5 μm s-1; e.g. søndergaard et al. 2015; hüssy et al. 2016). this approach is rapid, suppresses depth heterogeneity and avoids downhole elemental fractionation, as it ablates only to a depth of a few microns. the potential interfering effects of varying scan speeds, washout times and debris blanketing from the ablation are not yet studied, but this is intended in the near future. for ongoing otolith studies, we used a line-scan la-icpms approach to measure 325 cod otoliths (15–30 otoliths per day). the otoliths were embedded in epoxy resin and cut along the dorsoventral profile to expose the core and annuli. for streamlined, high-throughput analyses, custom-made epoxy mounts 7 × 0.8 × 0.8 cm large were cast, each holding 10 otoliths. a custom-made 10 cm2 sample holder with four mounts allows for analysis of 40 otoliths per day. for each otolith, a reflected light optical photo depicting a line from core to edge (fig. 1d) was imported as an overlay image into the laser ablation software. the image was fitted to the otolith position in the live image, using reference points to facilitate the exact positioning of the line scans. the icpms analysis was optimised for dry plasma conditions through continuous linear ablation of the nist 612 standard. the signal-to-noise ratios were maximised for the isotopic mass range from mg to ba, while opting for low element-oxide production levels by minimising the 254uo2/238u ratio. instrumental drift was minimised by following a standard–sample–standard analysis protocol, bracketing every sample analysis by line analyses of the nist-612 and nist614 glass standards (jochum et al. 2011), while the febs-1 otolith (sturgeon et al. 2005) and the macs-3 carbonate powder tablets (jochum et al. 2012) provided quality control of the nist-612 standard measurements. the averaged 2σ accuracy and uncertainty of the standards were typically <5% for element abundances >1–3 ppm. data processing was done with software iolite v. 2.5 (hellstrom et al. 2008; paton et al. 2011) using the trace elements is data reduction routine. calculation of abundances were based on 43ca isotope as the internal standard, assuming 38.3 wt% ca in all otoliths, comparable to the certified ca concentration reported for the febs-1 otolith standard (sturgeon et al. 2005). application of the analytical approach on baltic cod the main purpose of the analyses is to answer research questions important to the fisheries management in denmark and greenland. the specific approach presented here and modified versions of this method are used in several projects addressing a range of management questions relating to stock structure, migrations and age determination. here, we highlight some preliminary results of the most advanced project. the project tagging baltic cod (tabacod) is a joint baltic collaboration aiming to develop a new age-estimation method based on seasonal variations in element concentrations along a gradient from the core (birth) to the edge (death) of otoliths. fish age is one of the key variables in stock table 1. instrument operating conditions, data acquisition and processing parameters instrumentation: thermo-fisher scientific element 2 double focusing sf-icp-ms forward power: 1470 w cones: ni plasma gas: 16 l min-1 auxiliary gas (he): 0.85 l min-1 nebuliser gas (ar): 0.95 l min-1 new wave research nwr 213 solid state nd:yag laser ablation system laser wavelength: 213 nm laser fluence: ~ 9.5 j cm-2 spot size: 40 µm repetition rate: 10 hz scan speed: 5 µm s-1 data acquisition and processing: analyte isotopes: 25mg, 31p, 43ca, 44ca, 55mn, 65cu, 66zn, 88sr, 137ba sampling time, ms: 10 10 10 10 10 10 10 10 10 samples per peak: 10 10 10 10 10 10 10 10 10 acquisition: time resolved (continuous analysis) along transects mass resolution: 300 (low) oxide production rate tuned to ≤0.3% uo2 (254uo2/238u) single analysis duration and setup: 30 s blank, 2–20 min ablation (sample dependent), 30 s washout software for data reduction: iolite version 2.5 (paton et al. 2011; hellstrom et al. 2008) standards: internal standard isotope: 43ca external standardisation: nist-612 glass secondary standards: nist-614 glass, febs-1 otolith powder and macs-3 carbonate powder pressed as tablets 93 assessment and has traditionally been obtained by visual examination of otolith cross-sections, where seasonal fluctuations in growth are visible as optically contrasting growth zones much like the rings in cross-sections of trees. in recent years, this traditional method has failed to provide reliable age information, thus posing severe management problems for the eastern baltic cod stock. initial results from the tabacod project on the seasonality in the otolith chemical composition are presented as an example of the application of the la-icpms approach. all analysed otoliths were acquired from cod that were subjected to a mark-recapture experiment. a total of c. 500 cod specimens were captured, externally marked, injected with srcl2 and released again. when the cod were recaptured they were sent to dtu aqua for analysis. the srcl2 is incorporated into the otolith as it grows and acts as an internal timestamp. the chemical signals from timestamp to edge corresponds to the time the fish spent at sea between capture and recapture. combining information on how many days the fish had been at sea, what time of year it was released/recaptured and how much it had grown since tagging allows us to validate our hypotheses on seasonally varying element concentrations. concentrations of elements like mg and zn vary with season (hüssy et al. 2016). figure 2 depicts core-to-edge compositional profiles of mg, p, mn, cu, zn and sr from the otolith shown in fig. 1d, corresponding to the entire life of the fish. the red line in fig. 2 marks where the sr concentration dramatically increases, representing the srcl2-tagging event. the mg, p and mn concentrations show clear and similar variations throughout the otolith structure on a scale of tens to hundreds of ppm. zn and cu concentrations vary around our analytical resolution threshold of c. 1–3 ppm and do not show significant systematic variations. the la-icpms data are currently undergoing statistical analysis to quantitatively identify seasonal variations. however, some analysed elements indicate clear patterns resembling seasonality. if a seasonality in element concentration occurs, superimposing individual transects (like the ones shown in fig. 2) of all cod in one plot will result in a generic signal. if element signals are random in relation to time, no such signal will be evident. figure 3 shows p concentrations from all analysed otoliths, standardised by dividing each measured value by the mean profile p concentration to remove the effect of differences in average p levels between individuals. all profiles were centred at the sr peak, since all cod were tagged during the same season (april to may of the same year). the time scale on the x-axis is estimated, assuming linear growth within years, similar growth in individuals across the year, and that all specimens were tagged and released on the same date of a given year. although there are individual differences between fish, fig. 3 indicates a general seasonal variation in distance to core μm 25mg 31p 55mn 65cu 66zn 88sr 20 40 60 ppm 250 500 750 1000 0 5 10 15 20 0.0 0.4 0.8 0 1 2 3 4 5000 10000 0 1000 2000 3000 4000 fig. 2. trace-element concentration profiles in ppm of the cod otolith shown in fig. 1d. the x-axis indicates the concentration along the profile from 0 µm, when the fish was born, to 4300 µm, when it died. the red line marks the position of the sr peak induced by srcl2 injection. data from nielsen et al. (2018). as this cod was tagged in april, the sr timestamp corresponds to the coldest water temperatures experienced by the cod over a year, where mg and p concentrations are at a minimum. −400 0 400 800 −100 0 100 200 days around centered sr peak positions m ea n ce nt er ed 3 1 p p pm fr om a ll ot ol ith s fig. 3. variations of phosphorous concentration in all measured otoliths c. 200 days before and after the srcl2 tagging and release experiment. on the x-axis, measurements are centred on the injection-induced sr peak. on the y-axis, measurements have been centred on the mean p concentration of each otolith transect. the red line is a generalised additive model smoothed curve. 9494 p concentrations. the total number of minima occurring in the profiles from birth to death thus corresponds to the number of winters the fish has experienced and hence its age. other ongoing projects migration patterns of the kattegat cod: this project seeks to map migration patterns of cod captured in the kattegat, which were genetically identified to belong to the north sea or western baltic stock. elemental profiles of 400 cod captured along a geographic gradient covering the entire kattegat will be compared with baseline samples from adjacent areas. comparing the results to information about the otoliths’ annual growth zones can reveal at what time in its life the cod has migrated to and from the kattegat. stock structure in capelin: with partners from greenland’s fishing industry this project investigates stock structure, migration and natal homing (the return to a birthplace to reproduce) of capelin (mallotus villosus) from 18 areas along the coasts of south and west greenland. the aim of this project is to provide counsel on sustainable management of a species that plays a vital role in the marine food chain. final remarks a la-icpms approach for quantitative, high-throughput transect measurements of otoliths was successfully set up at geus. data from 325 otoliths are being thoroughly examined, and only an ‘appetizer’ of the data is presented here. the la-icpms approach is adaptable for most solid carbonate (e.g. bivalves) and phosphate (e.g. teeth, horn) materials showing cross-surface compositional variations. analyte isotopes include most major, minor and trace elements and acquisition parameters are easily optimised for the specific sample type, thus providing a rapid and extremely versatile in-situ analytical approach for comparable natural materials. acknowledgements we thank balticsea2020 for financial support and mojagan alaei, geus, for laboratory assistance. references degens, e.t., deuser, w.g. & haedrich, r.l. 1969: molecular structure and composition of fish otoliths. marine biology 2, 105–113. hellstrom, j., paton, c., woodhead, j. & hergt, j. 2008: iolite: software for spatially resolved la-(quad and mc) icpms analysis. mineralogical association of canada short course series 40, 343–348. hüssy, k., hinrichsen, h.h., fey, d.p., walther, y. & velasco, a. 2010: the use of otolith microstructure to estimate age in adult atlantic cod gadus morhua. journal of fish biology 76, 1640–1654. hüssy, k., gröger, j., heidemann, f., hinrichsen, h.h. & marohn, l. 2016: slave to the rhythm: seasonal signals in otolith microchemistry reveal age of eastern baltic cod (gadus morhua). ices journal of marine science: journal du conseil 73, 1019–1032. jochum, k.p. et al. 2011: determination of reference values for nist srm 610–617 glasses following iso guidelines. geostandards and geoanalytical research 35, 397–429. jochum, k.p., scholz, d., stoll, b., weis, u., wilson, s.a., yang, q., schwalb, a., börner, n., jacob, d.e. & andreae, m.o. 2012: accurate trace element analysis of speleothems and biogenic calcium carbonates by la-icp-ms. chemical geology 318–319, 31–44. limburg, k.e., wuenschel, m.j., hüssy, k., heimbrand, y. & samson, m. 2018: making the otolith magnesium chemical calendar-clock tick: plausible mechanism and empirical evidence. reviews in fisheries science & aquaculture 26, 1–15. nielsen, k., serre, s., thomsen, t. & hüssy, k. 2018: using la-icpms to investigate seasonality in cod otolith microchemistry. 33rd nordic geological winter meeting, dtu, kongens lyngby, denmark. abstract volume. pannella, g. 1971: fish otoliths: daily growth layers and periodical patterns. science 173, 1124–1127. paton, c., hellstrom, j., paul, b., woodhead, j. & hergt, j. 2011: iolite: freeware for the visualisation and processing of mass spectrometric data. journal of analytical atomic spectrometry 26, 2508–2518. patterson, w.p., smith, g.r. & lohmann, k.c. 1993: continental paleothermometry and seasonality using the isotopic composition of aragonitic otoliths of freshwater fishes. in: swart, p.k. et al. (eds): climate change in continental isotopic records. american geophysical union monograph (1993), 191–202. søndergaard, j., halden, n., bach, l., gustavson, k., sonne, c. & mosbech, a. 2015: otolith chemistry of common sculpins (myoxocephalus scorpius) in a mining polluted greenlandic fjord (black angel leadzinc mine, west greenland). water, air and soil pollution 226(10) 336, 12 pp. sturgeon, r.e., willie, s.n., yang, l., greenberg, r., spatz, r.o., chen, z., scriver, c., clancy, v., lam, j.w. & thorrold, s. 2005: certification of a fish otolith reference material in support of quality assurance for trace element analysis. journal of analytical atomic spectrometry 20, 1067–1071. sturrock, a.m., trueman, c.n., darnaude, a.m. & hunter, e. 2012: can otolith elemental chemistry retrospectively track migrations in fully marine fishes? journal of fish biology 81, 766–795. thomas, o.r.b., ganio, k., roberts, b.r. & swearer, s.e. 2017: trace element–protein interactions in endolymph from the inner ear of fish: implications for environmental reconstructions using fish otolith chemistry. metallomics 9, 239–249. thorrold, s.r. & shuttleworth, s. 2000: in situ analysis of trace elements and isotope ratios in fish otoliths using laser ablation sector field inductively coupled plasma mass spectrometry. canada journal of fisheries and aquatic sciences 57, 1232–1242. authors’ addresses s.h.s. & t.b.t., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k., denmark. e-mail: shs@geus.dk. k.e.n., p.f-j. & k.h., national institute of aquatic resources, technical university of denmark, kemitorvet, building 202, dk-2800 kgs. lyngby, denmark. mailto:shs@geus.dk geological survey of denmark and greenland bulletin 23, 2011, 57–60 57 the tartoq group is located in the sermiligaarsuk fjord region in south-west greenland in an area of approximately 20 × 50 km (fig. 1). the tartoq group consists of several discrete, fault-bound blocks of metavolcanic rocks, surrounded by archaean tonalite-trondhjemite-granodioritetype (ttg) gneisses. a zircon age of 2996.3 ± 5.9 ma of a ttg intrusion provides a minimum age for the formation of the tartoq group (fig. 2). the metavolcanic rocks probably show the lowest degree of metamorphism found anywhere in the archaean craton of greenland. here we present a new model for the origin of the metavolcanic rocks of the tartoq group based on geochemical, metamorphic and structural data. the samples used for this study were collected by the geological survey of denmark and greenland (geus) in 2009 and 2010. the study is part of a joint project between the greenland bureau of minerals and petroleum and geus on the mineral potential of south-western greenland. geology of the tartoq group the tartoq group can be divided into five main lithological units: (1) subaqueous, mafic pillow lavas, (2) mafic dykes and sills with semi-ophitic textures, (3) gabbros with relict magmatic textures and tonalite-trondhjemite-granodiorite composition, (4) ultramafic rocks that are mainly serpentinites and (5) felsic schists that commonly show high strain with mylonitic textures. in addition, a marble unit occurs in the nuuluk block; gneiss of quartz dioritic composition lies in a large breccia zone with intrusive tonalite-trondhjemitegranodiorite-type (ttg) gneisses that have been fractured and chloritised; and undifferentiated pegmatites are found in the bikuben and the iterlak blocks. the lithologies and their relations were described by higgins (1968), berthelsen & henriksen (1975), petersen (1992) and van hinsberg et al. (2010). greenschist facies rocks with pseudo-sections indicating peak metamorphic conditions of 380°c at 2 kbar dominate the nuuluk block in the west (fig. 1). the metamorphic grade increases to upper amphibolite facies with pressure and temperature calculated to 650°c at 6–7 kbar in the bikuben block in the east. evidence for partial melting is seen for at least one locality in the bikuben block, where leucosomes occur in fold hinges of amphibolites. peak metamorphic assemblages generally correlate with the different lithological units, which means that pillow lavas and shallow dykes or sills are at a lower grade and gabbro and serpentinite units at a higher grade (van hinsberg et al. 2010). the tartoq group is thrust (top to the se) and imbricated with younger ttg gneisses. kilometre-scale nappes, low-angle shear zones, and younger cataclastites formed as a remnants of mesoarchaean oceanic crust in the tartoq group, south-west greenland kristoffer szilas, vincent j. van hinsberg, alexander f. m. kisters, thomas f. kokfelt, anders scherstén and brian f. windley 48°w 49°w 61°45'n 61°40'n 61°35'n 61°30'n 61°25'n61°25´n 61°30´n 48°30´w49°w 48°w iterlak bikubennuuluk amitsuarsua naalagaaffik ser milig aa rsu k nuna qaqortoq se rm ilig aa rsu k b ræ midternæs akuliaruseq 5 km quaternary cover tartoq group supracrustal rocks thrust faults in study area felsic schist serpentinite chloritised breccia zone pegmatite, undifferentiated orthogneiss, undifferentiated ketilidian supracrustal rocks granite, undifferentiated 61°40´n 61°35´n 61°45´n archaean proterozoic greenland fig. 1. simplified geological map of the sermiligaarsuk fjord area, with the names of seven blocks with supracrustal rocks belonging to the tartoq group. the undifferentiated orthogneiss surrounding tgg gneisses has ages ranging from about 2500 ma to 3000 ma. the ketilidian supracrustals overly the tartoq group unconformably (van hinsberg et al. (2010). the tartoq group is mostly in tectonic contact with the surrounding orthogneiss, but a few intrusive relations are preserved at south iterlak and north amitsuarsua blocks. the supracrustal blocks form thrust sheets and lateral ramps with internal top to the se kinematic indicators. © geus, 2011. geological survey of denmark and greenland bulletin 23, 57–60. open access: www.geus.dk/publications/bull 5858 result of the accretion and progressive exhumation of these rocks. deformation was associated with hydrothermal alteration (with au mineralisation at c. 450°c) that overprinted regional metamorphic parageneses and intense carbonation in high strain zones. felsic, 1–80 m thick schist bodies in the mafic sequence represent high-strain zones where ttg gneisses have been preferentially sheared in the metavolcanic rocks. some of these mylonitic, felsic schists gave zircon ages of c. 2800 ma to 3700 ma which indicate the subsurface presence of old crustal material. these ages also suggest that the felsic schists are not part of the tartoq group, but are tectonic in origin. geochemistry the geochemical data were screened and samples showing evidence of post-magmatic alteration were rejected. microscopy of thin sections was used to identify the least altered samples. samples with quartz and carbonate veining and samples showing excessive large-ion lithophile elements (lile) enrichment were discarded. pillow lavas are prone to sea-floor alteration during extrusion, and therefore this group of rocks may be under-represented in the screened dataset. the tartoq group metavolcanic rocks (pillow lavas, dykes or sills and gabbro units) all have tholeiitic basaltic compositions. trace-element variations within the volcanic sequence broadly show that the incompatible trace elements increase with decreasing mgo content and the compatible trace elements decrease with falling mgo, consistent with trends of fractional crystallisation of olivine, clinopyroxene and perhaps plagioclase. the major elements show trends similar to what would be expected for a tholeiitic fractionation series. the scatter seen within the data may be due to metamorphic or slight hydrothermal background alteration. the metavolcanic rocks generally have flat, primitive, mantle-normalised trace-element patterns with lan/smn = 0.8–1.0, but they show negative nb anomalies with nbn/ lan = 0.4–0.9 (fig. 3). in tectonic discrimination diagrams, which rely on immobile trace elements, the metavolcanic rocks plot in the mid-ocean ridge basalt (morb) or islandarc tholeiite (iat) fields. la, y and nb abundances are similar to those of some modern back-arc basalts (bab). the serpentinites have median values of sio2 (45 wt%), mgo (37 wt%), feot (14.5 wt%), cr (3100 ppm), ni (600 ppm) and essentially no cao or al2o3, combined with a large negative chondrite-normalised eu anomaly (fig. 3). the serpentinites mainly consist of normative olivine and hypersthene (3:1 ratio) and are thus harzburgitic in composition. the felsic schist units show complete overlap in major and trace elements with the surrounding gneisses. this is consistent with the structural interpretation that they are deformed rocks along thrust faults, and hence unrelated to the volcanic sequence. discussion the metavolcanic rocks in the tartoq group have similar trace-element patterns, with pronounced negative nb anomalies (fig. 3). their overall major and trace-element variations are consistent with fractional crystallisation processes. these features suggest that the metavolcanic rocks are co-magmatic. the dykes and sills have the highest concentration of incompatible trace elements. this is consistent with the fact that they have the lowest mgo content and they thus form the more evolved portion of the magma pile. the four pillow lavas show a narrow compositional range that 6 4 2 0 n u m b er 2800 2850 2900 2950 3000 3050 3100 3150 3200 207pb/206pb age (ma) r elative p ro b ab ility a ge ( m a) 3080 3040 3000 2960 2920 2880 0.65 0.60 0.55 0.50 207pb/235u 2 0 6 p b /2 3 8 u 13 15 17 19 21 2750 2850 data-point error ellipses are 2σ 3150 data-point error symbols are 2σ fig. 2. zircon u/pb age of sample ggu 510771 using laser ablation inductively coupled mass spectrometry. the sample is an intrusive ortho gneiss from the northern contact in the amitsuarsua block and the age was 2996.3 ± 5.9 ma (95% confidence interval). the zircons in the sample are prismatic and show oscillatory growth zonation. only the zircon cores were analysed. spots that are concordant within ± 10% have been used and they form a tight normal distribution with a mean square weighted deviation of 0.76 and a probability of 0.85. the mean was weighted by data point errors only (no points were rejected). the analyses cut the concordia with minimal signs of lead loss and thus provide a robust age. 59 overlaps with the dykes and sills, as to be expected if the latter represent the feeding channels of lava flows. a broader compositional range is observed for the gabbros. positive eu anomalies, low trace element concentrations combined with high contents of cao, al2o3 and sr suggest some accumulation of plagioclase. the serpentinites show low incompatible trace-element concentrations, major element contents and normative compositions. these features are consistent with harzburgite that could represent either cumulates of the volcanic sequence or the residual source mantle. pge patterns show depletion in pt and pd which is also observed in peridotite xenoliths from kimberlites in the area (wittig et al. 2010). the serpentinites probably represent a sub-continental lithospheric mantle, which experienced high degrees of melt extraction and was exhumed together with the supracrustal rocks of the tartoq group. no co-genetic, calc-alkaline rocks have been found in the tartoq volcanic sequence and it is therefore unlikely that the volcanic rocks erupted through continental crust or formed in a mature arc setting. at first glance the geochemical features of the metavolcanic rocks resemble those of the modern morb, but they differ distinctly by having a negative nb anomaly and by their pattern of undepleted light rare-earth elements. the tectonic discrimination diagrams point to a iat setting, but the lack of co-genetic, calc-alkaline rocks argues against this. the la, y and nb concentrations indicate a bab setting, which is in agreement with the arc ‘flavour’ that is suggested by the iat affinity. the major and trace-element values for the tartoq group are similar to data from archaean tholeiitic rocks from the superior province of canada, which formed in a bab setting according to the interpretation of sandeman et al. (2006). another explanation for archaean tholeiites is that they formed in an oceanic plateau setting similar to the recent ontong java plateau (arndt et al. 1997). all examples of archaean tholeiites are associated with abundant komatiites. however, komatiites are rare in the supracrustal belts of greenland and absent in the tartoq group, which argues against an oceanic plateau setting. rocks from a bab setting are also more likely to be preserved over time, because the crust in bab settings is thin and located in a collision setting, like the present-day lau and mariana back-arc crust (martinez & taylor 2003). it is mainly the higher degree of partial melting estimated for archaean tholeiites compared to present day morb settings that have led some authors to argue for an oceanic plateau setting, whereas the lile enrichment and negative dykes and sills sa m p le / p ri m it iv e m an tl e 0.1 1 10 gabbros pillow lavas sa m p le / p ri m it iv e m an tl e nb la ce pr nd zr sm eu ti dy y yb lu nb la ce pr nd zr sm eu ti dy y yb lu 0.1 1 10 serpentinites fig. 3. spidergrams showing primitive mantlenormalised trace-element compositions of samples from the tartoq group (normalisation after sun & mcdonough 1989). only the relatively immobile trace elements are shown due to the likely mobility of the large ion lithophile elements during metamorphism and seafloor alteration. the pillow lavas, dykes and sills form fairly well-constrained compositional groups. in contrast, the gabbros show greater compositional variation, reflecting variable effects of fractionation and accumulation processes. the serpentinites show low concentrations and a large negative eu anomaly, which imply melt or mineral equilibrium processes involving plagioclase. the yellow areas shown in the spidergrams represent the total data field of the tartoq group. 6060 nb anomaly have led others to suggest a bab environment. however, rollinson (2010) suggested that the differences between modern morb and archaean non-arc tholeiites simply reflect high temperatures of the archaean mantle. rollinson’s model provides a robust explanation for the geochemical features of the tartoq group. the tartoq group may be a product of a hotter mantle giving rise to a thicker melting column, which affected the composition and differentiation of morb magmas and resulted in the observed discrepancies compared with the modern morb. regardless of the precise setting of ocean crust formation, the similar flat trace-element patterns and fractional crystallisation trends of all the metavolcanic rocks of the tartoq group, together with the presence of serpentinites, indicate that the rocks form a co-magmatic assemblage resembling an ophiolitic ocean floor sequence. two possible scenarios can explain the metamorphic and structural observations: (1) shallow subduction followed by ttg formation due to slab melting made the oceanic crust sufficiently buoyant to cause exhumation in a subduction channel and subsequent incorporation into the overriding plate; (2) the tartoq group formed from the overriding plate in an oceanic flake-style subduction setting that was dragged down with a shallow pt-trajectory by the subducting plate and later rebounded during aborted subduction (or slab break-off?), which could initiate ttg formation by decompression melting of the lower crust. in both scenarios we envisage that subduction took place in an intra-oceanic setting, resulting in the lack of continent-derived material in the tartoq group. both tectonic environments could give rise to the observed peak metamorphic assemblages and later retrogression by fluid input together with inter-thrusting with the ttg gneisses, which resulted in the tectonic slices and slabs that we see today. conclusions based on geochemical, metamorphic and structural data we have developed a new model according to which the tartoq group is a slab of oceanic crust. we interpret the protolith of the tartoq group as a structurally dismembered section of archaean oceanic crust of either morb or bab affinity, which might provide valuable insight into archaean geodynamics. references arndt, n.t., kerr, a.c. & tarney, j. 1997: dynamic melting in plume heads: the formation of gorgona komatiites and basalts. earth and planetary science letters 146, 289–301. berthelsen, a. & henriksen, n. 1975: geological map of greenland, 1:100 000, ivigtut 61 v.1 syd. descriptive text, 169 pp. copenhagen: geological survey of greenland. higgins, a.k. 1968: the tartoq group on nuna qaqortoq and in the iterdlak area, south-west greenland. rapport grønlands geologiske undersøgelse 17, 17 pp. martinez, f. & taylor, b. 2003: controls on back-arc crustal accretion: insights from the lau, manus and mariana basins. in: larter, r.d. & leat, p.t. (eds): intra-oceanic subduction systems: tectonic and magmatic processes. geological society special publications (london) 219, 19–54. petersen, j.s. 1992: nuuluk-iterlak gold and massive-sulfide project, taartoq archaean greenstone belt, sw greenland, 164 pp. unpublished field report, nunaoil a/s. rollinson, h. 2010: coupled evolution of archean continental crust and subcontinental lithospheric mantle. geology 38, 1083–1086. sandeman, h.a., hanmer, s., tella, s., armitage, a.a., davis, w.j. & ryan, j.j. 2006: petrogenesis of neoarchaean volcanic rocks of the macquoid supracrustal belt: a back-arc setting for the northwestern hearne subdomain, western churchill province, canada. precambrian research 144, 140–165. sun, s. & mcdonough, w. f. 1989: chemical and isotopic systematics of oceanic basalts: implications for mantle composition and processes. in: saunders, a.d. & norry, m.j. (eds): magmatism in the ocean basins. geological society special publications (london) 42, 313–345. van hinsberg, v.j., szilas, k. & kisters, a.f.m. 2010: the tartoq group, sw greenland: mineralogy, textures and a preliminary metamorphic to hydrothermal history. danmarks og grønlands geologiske undersøgelse rapport 2010/120, 40 pp. wittig, n., webb, m., pearson, d.g., dale, c.w., ottley, c.j., hutchinson, m., jensen, s.m. & luguet, a. 2010: formation of the north atlantic craton: timing and mechanisms constrained from re–os isotope and pge data of peridotite xenoliths from s.w. greenland. chemical geology 276, 166–187. authors’ addresses k.s. & t.f.k., geological survey of denmark and greenland, øster voldgade 10, 1350 copenhagen k, denmark. e-mail: ksz@geus.dk v.j.h., department of earth sciences, university of oxford, south parks road, oxford ox1 3an, uk. a.f.m.k., department of earth sciences, stellenbosch university, matieland 7602, south africa. a.s., deptartment of earth and ecosystem sciences division of geology, lund university sölvegatan 12, 223 62 lund, sweden. b.f.w., department of geology, university of leicester, university road, leicester, le1 7rh, uk. geological survey of denmark and greenland bulletin 33, 2015, 77-80 77 investigations of detrital zircon, rutile and titanite from present-day labrador drainage basins: fingerprinting the grenvillean front tonny b. thomsen, christian knudsen and alana m. hinchey a multidisciplinary provenance study was conducted on stream sediment samples from major rivers in the eastern part of labrador, canada (fig. 1). th e purpose was to fi ngerprint the sources that deliver material to the stream sediments and to the reservoir sand units deposited off shore in the sedimentary basins in the labrador sea. we used a multimineral u-pb geochronological approach employing rutile and titanite in addition to zircon to obtain unbiased age data. th e purpose of this was to characterise the diff erent igneous and metamorphic episodes that occurred in labrador, which is an area with highly variable geology characterised by the palaeoproterozoic south-eastern churchill province in the north-west, the archaean nain plutonic suite in the north-east, the palaeoproterozoic makkovik province in the east and the mesoproterozoic grenville province to the south. th e fi eld work was carried out in 2012 and 2013 and the study is a collaborative project between the geological survey of denmark and greenland and the geological survey of newfoundland and labrador. in this paper we focus on three samples from the southern part of the study area where two parts of the grenville orogeny are found (fig. 1). the use of zircon, rutile and titanite in sedimentary provenance investigations methods for obtaining geochronological information from various detrital minerals for quantitative sedimentary provenance purposes have developed rapidly over recent years. th is is mostly due to advances in high-throughput microanalytical techniques such as laser ablation inductively coupled plasma mass spectrometry (la-icpms). th e emphasis on u-pb geochronology has primarily been on detrital zircon, as it typically provides precise age information of the source rocks. zircon, however, is not usually the mineral of choice for dating the history of rocks with a complex tectonothermal evolution, as it typically survives most processes occurring in the rock cycle from sedimentation to high grade metamorphism and oft en even magmatic processes. th is means that zircon typically refl ects several orogenic cycles (e.g. okay et al. 2011) and is less suited for recording information about processes related to the metamorphic or hydrothermal reaction history of a rock. th e ti minerals, rutile and titanite, occur in a variety of magmatic, hydrothermal, metamorphic and sedimentary rock types, oft en together with zircon (e.g. force 1991; frost et al. 2001; zack et al. 2004). although detrital rutile is less abundant than zircon, the widespread occurrence of rutile in a wide range of mediumto high-grade, and also high-pressure (blueschist and eclogite facies), metamorphic rocks as well as in sediments and sedimentary rocks (force 1980, 1991), combined with its high mechanical and chemical stability during weathering, transport and diagenesis (e.g. morton & hallsworth 1999), makes it a prime candidate in provenance studies (zack et al. 2011). rutile forms under amphibolite and higher metamorphic facies conditions and is typically unstable at lower grade conditions (force 1980, © 2015 geus. geological survey of denmark and greenland bulletin 33, 77–80. open access: www.geus.dk/publications/bull ggu 539754 ggu 539828 ggu 539845 proterozoic nain plutonic suite grenville province exterior thrust belt interior magmatic belt makkovik and se churchill provinces archaean superior province nain province sample drainage area 100 km 54°n 52°n 54°n 60°n 66°w 62°w 56°w 62°w66°w fig. 1. simplified geological map of labrador. 7878 1991; triebold et al. 2007, 2011), where it usually breaks down to form other ti minerals such as titanite or ilmenite at greenschist facies conditions. rutile, therefore, typically yields chronological and petrogenetic information refl ecting the timing and conditions of the last medium to high-grade or high-pressure metamorphic event (zack et al. 2011; okay et al. 2011). titanite is widespread in a variety of rock types typically of more calcic compositions, and is usually rare in rocks with low cao/al2o3 ratios such as peralkaline granitoids and peraluminous granites (frost et al. 2001). titanite occurs in very lowto high-grade metamorphic rocks and survives under ultra-high pressure conditions (force 1991; frost et al. 2001), although it is typically scarce or absent in most granulite-facies metamorphic rocks (krogh & keppie 1990). even though titanite is a widespread mineral and occurs as detrital and authigenic grains in sedimentary rocks, it has rarely been used to date deposition, diagenesis or low-grade metamorphism. detrital titanite was fi rst used as a sedimentary provenance tool by mcateer et al. (2010). th e reason that titanite is not routinely used for provenance investigations is probably that it is more susceptible to abrasion during sedimentary transport than zircon and rutile, resulting in reduced occurrence or absence as a detrital component in sediments and sedimentary rocks. titanite, however, is more reactive than zircon or rutile during metamorphism and forms at temperatures below 700°c. th is provides the mineral with a large potential to record ages for a wide range of low to moderate temperature geological crustal processes (frost et al. 2001; mcateer et al. 2010; muhling et al. 2012). th us, if present in sedimentary rocks, titanite is an ideal candidate for dating regional or local metamorphism. in addition, because most metamorphic events are associated with deformation, titanite can date potential deformation stages in metamorphic terranes (frost et al. 2001). th erefore, it is possible to recognise sediment sources from detrital titanite and rutile data that are not represented in zircon data and thus gain additional chronological and petrogenetic insight into the tectonothermal history of the source regions (mcateer et al. 2010, 2014). furthermore, rutile and titanite generally contain 5–10 times less u than zircon, thus metamictisation of these minerals is relatively rare. consequently, in rocks with u-rich zircon and titanite or rutile, the zircon might show metamictisation, and therefore would be more prone to degradation during transport and weathering, and is thus likely to be excluded from a detrital study (fedo et al. 2003). titanite and rutile, on the other hand, have the potential to retain the magmatic record of the u-rich zircon source as well as the metamorphic episodes that may have occurred prior to deposition. u-pb analysis and data processing in this study, u-pb dating was carried out on mineral grains embedded in epoxy mounts at the la-icpms facility at the geological survey of denmark and greenland using a nwr213 nd:yag laser system coupled to an element 2 double-focusing, single-collector, magnetic sector-fi eld icpms. mineral grains were separated by routine separation methods including a wilfrey water-shaking table, frantz electromagnetic separation and heavy liquids. zircon, rutile and titanite grains were picked by hand under a binocular microscope from the resulting heavy mineral fractions and mineral compositions were qualitatively controlled by semeds. for rutile and titanite, laser beam pre-ablation using a spot size slightly larger (40 μm) than the analysis spot size (25 μm) was performed prior to the la-icpms analysis to avoid surface contamination. data processing was performed off -line using the soft ware iolite (paton et al. 2010, 2011) and the vizualage data reduction scheme vers. 2.5 by petrus & kamber (2012). th e data were corrected for background signal, time-dependent fractionation, instrumental drift and down-hole isotopic fractionation. in order to validate our results, the measurements were bracketed throughout the entire analysis sequences by analyses of natural mineral standards. th ese include the gj-1 and plesovice zircons, the rutiles r10, r13, r19 (provided by courtesy of t. zack, university of gothenburg) and sugluk-4 (l. bracciali, british geological survey), and the titanite a1772 (y. lahaye, geological survey of finland) and seiland (j. kosler, university of bergen). in contrast to zircon, common pb in titanite and rutile is generally lattice bound and can occur in signifi cant proportions. th us, common pb correction typically needs to be applied for these minerals. however, common pb usually has negligible eff ects for titanites or rutiles with 206pb/204pb ratios >300 (frost et al. 2001), whereas the initial common pb isotopic composition typically has greater eff ects on analyses with 206pb/204pb ratios <300 and therefore the results must be treated more cautiously. only the titanite grains in this study typically have large proportions of common pb, whereas most zircon and rutile only contain a small or negliable amount of common pb. titanite ages reported herein are corrected for common pb. some of the titanite grains contained a signifi cant amount of common pb and some of the titanite ages are potentially of a slightly lower accuracy compared to ages obtained for rutile and zircon that were not common pb corrected. th e correction for common pb was performed using the present-day terrestrial common pb estimate of stacey & kramers (1975) and the measured mass 204 (204pb + 204hg) corrected for 204hg calculated from measured 202hg and the natural 204hg /202hg ratio. 79 results of the provenance study u-pb age distributions of zircon, rutile and titanite from three representative river samples are shown in the probability–density diagrams in fig. 2. th e three samples were collected in areas within the grenvillean orogenic belt that are characterised by ages around 1000 ma. sample ggu 539754 comes from the northern part of the orogenic belt (the exterior thrust belt), sample ggu 539828 from the central part and sample ggu 539845 from the southern part (the interior magmatic belt). th e samples all contain zircon grains older than the grenville orogeny, refl ecting the ability of zircon to maintain older magmatic formation signatures through the younger grenvillean orogenesis. th e three areas show distinct diff erences in detrital zircon ages. th e frequency of c. 1000 ma old grenvillean zircon ages is much lower in the northern exterior thrust belt than in the central part and especially in the southern interior magmatic belt, where crust was formed during the grenville orogeny. in the southern area (ggu 539845) both palaeoand mesoproterozoic zircon ages are common, with the highest abundance at c. 1500 ma, whereas the samples from the central area (ggu 539828) and the northern exterior thrust belt (ggu 539754) are dominated by palaeoproterozoic zircons with a peak at c. 1650 ma. however, there is a second distinct mesoproterozoic peak at c. 1360 ma in the northern area and a signifi cantly reduced abundance of mesoproterozoic zircon ages in the central area. as expected, the detrital rutile ages peak just below 1000 ma for all three samples, refl ecting rutile formation or complete u-pb system resetting of older rutile grains during mediumto high-grade metamorphic stages of the grenville orogeny. titanite shows grenvillean ages for all three samples. in the sample from the southern interior magmatic belt all titanite ages are c. 1000 ma old, corresponding to the rutile age distribution, whereas a more complex age pattern, with palaeoand mesoproterozoic titanites, is seen in the samples from the central and northern areas. th is indicates that many more titanite grains from the latter areas survived the grenvillean metamorphism than (1) rutile from the same areas and, (2) titanite and rutile from the southern interior magmatic belt. th e occurrence of pre-grenvillean titanite and absence of pre-grenvillean rutile in the central and northern areas could be due to the diff erence in the u-pb isotopic system closure temperatures of the two minerals; rutile has a lower closure temperature at c. 400–500°c than titanite with a closure temperature at c. 500–700°c. hence the age pattern depends on the metamorphic grade to which minerals were exposed during the grenville orogeny. th e occurrence of titanite ages and lack of rutile ages indicate mediumto high-grade metamorphic conditions, probably upper amphibolite facies. during the grenville orogeny, the metamorphic grade of the new crust was probably higher in the southern interior magmatic belt than in the central and northern areas. we suggest that the closure temperature of n = 10/16 n = 5/11 n = 32/32 n = 4/4 n = 140/140 n = 129/130 n = 8/9 n = 26/26 n = 136/137 0 0.00108 0.00217 0.00325 0.00434 0.00542 0 0.00245 0.00491 0.00736 0.00982 0.01227 p ro b ab ili ty 0 0.00214 0.00428 0.00642 0.00857 0.01071 0 0.00266 0.00533 0.00799 0.01066 0.01332 p ro b ab ili ty 0 0.00119 0.00238 0.00357 0.00477 0.00596 0.00076 0.00152 0.00228 0.00304 0.0038 p ro b ab ili ty 12 10 8 6 4 2 0 12 10 8 6 4 2 0 12 10 8 6 4 2 0 2 1 0 2 1 0 2 1 0 fr eq u en cy 0.00149 0 0.00299 0.00448 0.00598 0.00747 fr eq u en cy 0 0.0023 0.0046 0.00689 0.00919 0.01149 0 0.00118 0.00236 0.00354 0.00472 0.0059 fr eq u en cy 0 500 1000 1500 2000 t it an it e r u ti le z ir co n age (ma) 0 500 1000 1500 2000 age (ma) 0 500 1000 1500 2000 age (ma) ggu 539754 the nortern exterior thrust belt ggu 539828 the central part ggu 539845 the southern interior magmatic belt 30 25 20 15 10 5 0 30 25 20 15 10 5 0 30 25 20 15 10 5 0 0 fig. 2. probability–density diagrams showing u–pb age distributions for three samples collected in labrador. n = x/y (in red) denotes the total number of analysed zircon grains (y) of which x are ‘concordant’ (i.e. <10% discordant from concordia). light grey: ages that are >10 % discordant (relative to wetherill concordia), dark grey: ‘concordant’ ages within <10% discordance (i.e. <10%). blue line at 1000 ma: late stage of the grenville orogeny. 8080 both ti minerals was overstepped, resulting in isotopic age resetting of any pre-grenvillean titanite and rutile grains and formation of new grenvillean age titanite and rutile. concluding remarks th e wide range in the detrital zircon u-pb ages within the grenvillean orogenic belt refl ects formation age diff erences of the rocks that were brought into the orogenic process. accordingly, for the best characterisation of the sediment source, it is not suffi cient to know the age of the orogeny that aff ected the area; it is also nessesary to know the lateral distribution of the rock units in the area as well as the ages of the zircons (and other minerals) from these units. moreover, diff erent units contain diff erent amounts of e.g. zircon. an effi cient way to map the lateral age variation is to analyse detrital minerals collected from present-day drainage systems. references fedo, c.m., sircombe, k.n. & rainbird, r.h. 2003: detrital zircon analysis of the sedimentary record. in: hanchar, j.m. & hoskin, p.o. (eds): zircon: experiments, isotopes and trace element investigations. mineralogical society of america, reviews in mineralogy 53, 277–303. force, e.r. 1980: th e provenance of rutile. journal of sedimentary petrology 50, 485–488. force, e.r. 1991: geology of titanium-mineral deposits. geological society of america, special papers 259, 112 pp. frost, b.r., chamberlain, k.r. & schumacher, j.c. 2001: sphene (titanite): phase relations and role as a geochronometer. chemical geology 172, 131–148. krogh, t.e. & keppie, j.d. 1990: age of detrital zircon and titanite in the meguma group, southern nova scotia, canada: clues to the origin of the meguma terrane. tectonophysics 177, 307–323. mcateer, c.a., daly, j.s., flowerdew, m.j., connelly, j.n., housh, t.b. & whitehouse, m.j. 2010: detrital zircon, detrital titanite and igneous clast u–pb geochronology and basement–cover relationships of the colonsay group, sw scotland: laurentian provenance and correlation with the neoproterozoic dalradian supergroup. precambrian research 181, 21–42. mcateer, c.a., daly, j.s., flowerdew, m.j., whitehouse, m.j. & monaghan, n.m. 2014: sedimentary provenance, age and possible correlation of the iona group, sw scotland. scottish journal of geology 50, 143–158. morton, a.c. & hallsworth, c.r. 1999: processes controlling the composition of heavy mineral assemblages in sandstones. sedimentary geology 124, 3–29. muhling, j.r., rasmussen, b. & fletcher, i.r. 2013: dating deposition and low-grade metamorphism by in situ u–pb geochronology of titanite. mineralogical magazine 77, 1800 only. okay, n., zack, t., okay, a.i. & barth, m. 2011: sinistral transport along the trans-european suture zone: detrital zircon–rutile geochronology and sandstone petrography from the carboniferous fl ysch of the pontides. geological magazine 148, 380–403. paton, c., woodhead, j.d., hellstrom, j.c., hergt, j.m., greig, a. & maas, r. 2010: improved laser ablation u–pb zircon geochronology through robust downhole fractionation correction. geochemistry, geophysics, geosystems 11, 1–36. paton, c., hellstrom, j.c., paul, b., woodhead, j.d. & hergt, j.m. 2011: iolite: freeware for the visualisation and processing of mass spectrometric data. journal of analytical atomic spectrometry 26, 2508–2518. petrus, j.a. & kamber, b.s. 2012: vizualage: a novel approach to laser ablation icp-ms u-pb geochronology data reduction. geostandards and geoanalytical research 36, 247–270. stacey, j.s. & kramers, j.d. 1975: approximation of terrestrial lead isotope evolution by a two-stage model. earth and planetary science letters 26, 207–221. triebold, s., von eynatten, h., luvizotto, g.l. & zack, t. 2007: deducing source rock lithology from detrital rutile geochemistry: an example from the erzgebirge, germany. chemical geology 244, 421–436. triebold, s., luvizotto, g., tolosana-delgado, r., zack, t. & von eynatten, h. 2011: discrimination of tio2 polymorphs in sedimentary and metamorphic rocks. contributions to mineralogy and petrology 161, 581–596. zack, t., von eynatten, h. & kronz, a. 2004: rutile geochemistry and its potential use in quantitative provenance studies. sedimentary geology 171, 37–58. zack, t., stockli, d.f., luvizotto, g.l., barth, m.g., belousova, e., wolfe, m.r. & hinton, r.w. 2011: in situ u-pb rutile dating by la– icp–ms: 208pb correction and prospects for geological applications. contributions to mineralogy and petrology 162, 515–530. authors’ adresses t.b.t. & c.k., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: tbt@geus.dk a.m.h., geological survey of newfoundland and labrador, 50 elizabeth avenue, st. john’s, nl, canada. geological survey of denmark and greenland bulletin 7, 2004, p 45-48 45 the present-day storebælt (great belt), the waterway between the islands of fyn and sjælland (fig. 1), contains deeply incised valleys, locally more than 50 m deep, and is of crucial importance to the water exchange between the fully marine kattegat and the brackish baltic sea. the role of this important gateway changed significantly during the late and post-glacial period (since 15 000 b.p.), when the baltic basin experienced alternating freshwater, brackish and marine conditions as a result of changes in relative sea level (figs 2, 3). the importance of the storebælt in understanding the dynamics of the baltic basin is reflected in the large number of studies carried out (see bennike et al. 2004). the first detailed sedimentological and stratigraphic studies in the storebælt area that demonstrated the presence of early holocene freshwater deposits below the seabed were those of krog (1960, 1965, 1971), who also presented the first shore-displacement curve for the area (krog 1979). the balkat project the late and post-glacial evolution of the south-western baltic sea has been studied in detail during the past 15 years as part of the multi-disciplinary balkat project, a co-operation between the geological survey of denmark and greenland (geus), the baltic sea research institute in warnemünde, and other partners (jensen et al. 2002). acquisition and interpretation of shallow seismic data and sampling of vibrocores form the basis for sequence stratigraphic and sedimentological studies, which together with microand macropalaeontological studies have resulted in detailed interpretations of depositional environments. the chronology has been established by numerous radiocarbon dates. initial studies focused on the fakse bugt and gedser rev region and the history of the baltic ice lake (jensen & stecker 1992; lemke & kuijpers 1995; jensen et al. 1997) and were followed by detailed studies of the ancylus lake stage (jensen et al. 1999). further studies of relative shorelevel changes in the region (bennike & jensen 1998) and the femer bælt threshold (lemke et al. 2001) showed that local ice lakes developed in front of the retreating fennoscandian ice sheet. two major transgressions of the baltic ice lake are recorded, with maximum highstand levels of approximately 30 m and 20 m below present sea level (figs 2, 3; björck 1995). the ancylus lake transgression reached just above the southern threshold of the storebælt, and was followed by a regression that presumably exposed large parts of the former lake bottom (figs 2, 3). the identification of these lake stages in the femer bælt area show a much wider distribution than previously expected, with possible connections to the kattegat (yoldia sea, littorina sea) via the storebælt (fig. 3). studies in southern kattegat and the northern storebælt region have revealed that a late pleistocene relative sea-level the storebælt gateway to the baltic jørn bo jensen, ole bennike,wolfram lemke and antoon kuijpers geological survey of denmark and greenland bulletin 7, 45–48 (2005) © geus, 2005 falster langelandsbêlt germany sweden poland fakse bugt ø resund 12°e 14°e 54°n 55°n 10°e 56°n 20 m 20 m 40 m 20 m lille bælt g edser rev femer bælt baltic sea kattegat 50 km storebæ lt jylland fyn sjælland denmark fig. 1. present-day general bathymetry of the south-western baltic sea. location of fig. 5 (red frame) and the northern and southern thresholds in storebælt (red circles) are shown. highstand was followed by a lowstand during the latest pleistocene (12 000 – 11 500 b.p.; fig. 2). this was in turn succeeded by the littorina sea transgression, which resulted in a series of back-stepping coastal deposits (jensen et al. 2002). recent studies under the balkat project have been concentrated in the central storebælt area and the northern and southern threshold areas in order to obtain a more detailed understanding of the interaction between the kattegat and the baltic basin. the storebælt gateway the central storebælt area, between the fully marine kattegat and the brackish baltic basin, has been influenced by drainage from lakes in the baltic area and marine transgressions from the kattegat. both depend on the relative levels of the southern and northern thresholds. in this area the incised valley fills provide a unique opportunity to study the initial effects of drainage and transgression, as well as the timing of these events, that – with some delay – had a great influence on the baltic area. during the later stages of the balkat project, shallow seismic data and vibrocores have been collected for the area extending from the northern entrance to the southernmost part of the storebælt. the relative sea level changes in the southernmost kattegat region (fig. 2) are clearly recognised at the entrance to the storebælt, where late glacial marine highstand sediments are cut by the younger dryas lowstand erosional unconformity and followed by an early holocene succession of channel fill, estuary river mouth sediments and backstepping shoreface deposits (fig. 4a, b; bennike et al. 2000; jensen et al. 2002). the northern storebælt threshold is located in a less than 1 km wide incised valley (figs 1, 5). profiles north and south of the threshold (fig. 4c, d) show that marine transgressive deposits are found in the incised valley north of the threshold, whereas a transitional brackish unit exists below the littorina sea deposits south of the threshold. radiocarbon datings of the brackish sediments using terrestrial plant macrofossils indicate that the initial transgression of the littorina sea took place at about 9400 b.p. the central storebælt incised valley (figs 1, 4e) was formed by meltwater during the deglaciation about 17 000 b.p. (bennike et al. 2004), and the initial fill is represented by late glacial lake sediments. the youngest late glacial unit is restricted to the channels, and is believed to be the baltic ice lake extension into the storebælt area. the late glacial sediments are truncated by an erosional unconformity overlain by lower holocene freshwater sediments that include river and lakeshore deposits, and followed by extensive lake deposits formed in the time interval between 10 900 and 8800 b.p. (bennike et al. 2004). deposition of the early river deposits is coeval with the maximum level of the ancylus lake. the initial sign of the marine transgression in the central storebælt area is dated to 8100 b.p. by marine shells. 46 northern threshold ancylus lake baltic ice lake southern threshold littorina transgression 10 30 50 –10 –30 –50 sh o re -l ev el ( m ) years b.p. 16 000 12 000 8000 4000 0 central kattegat northern storebælt femer bælt fig. 2. shore level changes relative to present-day sea level in central kattegat, northern storebælt and femer bælt. the levels of the thresholds in the storebælt are indicated. 10 500 b.p.15 000 b.p. yoldia sea littorina sea 100 km 100 km baltic ice lake ancylus lake ice fig. 3. palaeogeographical maps showing the distribution of land and sea/lake at 15000 years and 10 500 years b.p. the threshold in the southernmost part of the storebælt is found in a few hundred metres wide channel at about 25 m below present sea level (figs 1, 4f, 5). however, fine-grained freshwater sediments dated to the time of maximum ancylus lake transgression at 10 300 b.p. brings the pre-ancylus lake threshold down to about 30 m below present sea level. there is no evidence of a rapid ancylus lake drainage (dana river) as earlier proposed by björck (1995). initial marine transgression of the storebælt based on the recently collected data it is possible to reconstruct a palaeogeographical scenario for the initial holocene marine transgression (10 000 – 9500 b.p.) of the storebælt area (fig. 5). at about 10 000 b.p. the ancylus lake was mainly drained by a river system located in the storebælt area with an outlet in the southern kattegat area (bennike et al. 2000). in general, the drainage pathway through the storebælt was restricted to channels less than 1 km wide. a transitional brackish estuary was restricted to the area immediately north-east of the northern threshold. the transgression of the littorina sea resulted in flooding of the northern storebælt threshold at about 9500 b.p. and a brackish environment extended to about 20 km south of the threshold. at the same time, a large local lake developed in the central and southern part of the storebælt area due to a ground-water level rise, related to the relative sea level rise north of the northern storebælt threshold, and brackish and marine conditions were gradually established in these areas around 9400–9100 b.p. (winn et al. 1998; bennike et al. 2004). future perspectives during the 15 years of balkat co-operation a unique database covering the late and postglacial sediments in the southwestern baltic region has been generated. the main key areas have now been studied, but further biostratigraphical investigations and datings are required for lillebælt, øresund and parts of kattegat. when this work is completed, a detailed model of the postglacial evolution of the western baltic region can be developed including data on palaeogeography, fauna and flora evolution, and climatic changes. this model will be of great importance for future scientific co-operation involving marine geological, archaeological, ecological and palaeo-climatic studies in the rest of the baltic region. furthermore a detailed knowledge on the palaeogeographic evolution is important in locating potential sand and gravel resources. large offshore construction works may also 47 sw ne –20 m –40 m –40 m –50 m –20 m –30 m –20 m –40 m –20 m –20 m –30 m –40 m 5 kmb ew sw ne lg w e lg lg gas w e ew 1km northern threshold southern threshold a c d e f gas 1km 1km 1km 1km till local ice lake baltic ice lake freshwater brackish marine highstand marine transgression littorina sea p le is to ce n e h o lo ce n e fig. 4. profiles illustrating the influence of the northern and southern thresholds in the zone between the fully marine kattegat and the brackish to lacustrine central and southern parts of the storebælt. for location of profiles, see fig. 5. benefit from the studies. for example, the planned femer bælt link bridge requires seabed information for geotechnical, raw material and hydrographic evaluations, as well as for monitoring possible impacts on the environment. acknowledgement our friend and colleague, wolfram lemke, unexpectedly passed away on 21 april 2005. we wish to acknowledge his enthusiastic participation in our joint projects and his inspiring contributions to our long-standing co-operation. his premature death is a great loss to the scientific community. references bennike, o. & jensen, j.b. 1998: lateand postglacial shore level changes in the southwestern baltic sea. bulletin of the geological society of denmark 45, 27–38. bennike, o., jensen, j.b., konradi, p.b., lemke, w. & heinemeier, j. 2000: early holocene drowned lagoonal deposits from the kattegat, southern scandinavia. boreas 29, 272–286. bennike, o., jensen, j.b., lemke, w., kuijpers, a. & lomholt, s.j. 2004: lateand postglacial history of the great belt, denmark. boreas 33, 18–33. björck, s. 1995: a review of the history of the baltic sea, 13.0–8.0 ka bp. quaternary international 27, 19–40. jensen, j.b. & stecher, o. 1992: paraglacial barrier–lagoon development in the late pleistocene baltic ice lake, southwestern baltic. marine geology 107, 81–101. jensen, j.b., bennike, o., witkowski, a. & kuijpers, a. 1997: the baltic ice lake in the south-western baltic: sequence-, chronoand biostratigraphy. boreas 26, 217–236. jensen, j.b., bennike, o., witkowski, a., lemke, w. & kuijpers, a. 1999: early holocene history of the southwestern baltic sea: the ancylus lake stage. boreas 28, 437–453. jensen, j.b., petersen, k.s., konradi, p., kuijpers, a., bennike, o., lemke, w. & endler, r. 2002: neotectonics, sea-level changes and biological evolution in the fennoscandian border zone of the southern kattegat sea. boreas 31, 133–150. krog, h. 1960: post-glacial submergence of the great belt dated by pollen-analysis and radiocarbon. report of the international geological congress, xxi session, part iv, 127–133. krog, h. 1965: on the post-glacial development of the great belt. baltica 2, 47–60. krog, h. 1971: the early post-glacial development of the storebælt as reflected in a former fresh water basin. quaternaria 14, 85–92. krog, h. 1979: the quaternary history of the baltic, denmark. in: gudelis, v. & königsson, l.-k. (eds): the quaternary history of the baltic, 207–217. uppsala: uppsala university. lemke, w. & kuijpers, a. 1995: late pleistocene and early holocene palaeogeography of the darss sill area, southwestern baltic. quaternary international 27, 73–81. lemke, w., jensen, j.b., bennike, o., endler, r., witkowski, a. & kuijpers, a. 2001: hydrographic thresholds in the western baltic sea: late quaternary geology and the dana river concept. marine geology 176, 191–201. winn, k., erlenkeuser, h., nordberg, k. & gustafsson, m. 1998: paleohydrography of the great belt, denmark, during the littorina transgression: the isotope signal. meyniana 50, 237–251. 48 authors’ addresses j.b.j, o.b & a.k., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: jbj@geus.dk w.l., baltic sea research institute, seestrasse 15, d-18119 rostock-warnemünde, germany. 10 000 b.p. f e d c a northern threshold southern threshold b 56°n 55°n 9 500 b.p. f e d c a northern threshold southern threshold b 56°n 55°n 11°e 11°e 25 km 25 km a b fig. 5. palaeogeographical maps of the storebælt area during initial holocene transgression. a: at c. 10 000 years b.p., before the transgression of the northern threshold, the ancylus lake was connected to kattegat via a river system ending in a narrow, brackish estuary. b: after the threshold was transgressed at c. 9000 years b.p., a much more extensive brackish water estuarine complex developed at the mouth of the river, and a major lake was formed in the central part of the storebælt area. for legend and sections a–f see fig. 4. geological survey of denmark and greenland bulletin 1, 777-811 777 the middle jurassic sortehat formation of the east greenland rift basin is a distinctive and laterally persistent mudstone-dominated succession sandwiched between the sandstone-dominated ostreaelv formation beneath and the sandstones of the pelion formation above (figs 1, 2). the stratigraphic terminology adapted palynostratigraphy and palaeoenvironment of the middle jurassic sortehat formation (neill klinter group), jameson land, east greenland eva b. koppelhus and carina f. hansen the grey–black mudstones of the sortehat formation form part of the middle jurassic fill of the jameson land basin in east greenland. the formation is exposed in the southernmost part of the north–south-trending, mesozoic rift system in east greenland that was part of the epeiric seaway between east greenland and norway. sedimentological observations of the sortehat formation indicate deposition in an offshore marine setting that was typically low energy and periodically oxygen-deficient but was influenced by storm currents on occasion. detailed palynological studies of the sortehat formation have resulted in the definition of three palynological assemblage zones recognised at four localities, namely enhjørningen dal and pelion (north jameson land), the type section at sortehat (central jameson land) and albuen at neill klinter along hurry inlet (south-east jameson land). in stratigraphic order, these zones are termed the botryococcus assemblage zone, the nannoceratopsis gracilis – nannoceratopsis senex assemblage zone, and the sentusidinium pelionense assemblage zone. they are recognised on the basis of the identification of approximately 110 species of palynomorphs, including 45 species of spores, 30 of pollen, 22 of dinoflagellate cysts, 10 acritarch species, two species of algae, and some fungal spores. an aalenian – ?early bajocian age is suggested for the sortehat formation on the basis of the palynoflora. interpretation of the palynomorph assemblages suggests that the formation accumulated in a shallow, brackish marine environment. a significant terrestrial input, including the freshwater green alga botryococcus, is recorded in the lower part of the formation and interpreted as an allochthonous accumulation in an offshore marine environment related to transgression of a low-lying coastal plain. a marked shift in the palynomorph assemblage seen by diversification of marine microplankton above the base of the formation, indicates an increase in the marine signal probably related to the onset of highstand conditions following the marine transgression. keywords: east greenland, jameson land basin, middle jurassic, aalenian – ?early bajocian, palynostratigraphy, sedimentology, palaeoenvironment, transgressive–highstand mudstones, allochthonous botryococcus assemblage e.b.k., geological survey of denmark and greenland, geocenter copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. present address: royal tyrrell museum of palaeontology, box 7500, drumheller t0j 0y0, alberta, canada. e-mail: evakoppelhus@hotmail.com c.f.h., geological institute, university of copenhagen, geocenter copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. present address: skovhegnet 4, dk-3460 birkerød, denmark. geological survey of denmark and greenland bulletin 1, 777–811 (2003) © geus, 2003 in this paper follows the revised scheme of dam & surlyk (1998). although the biostratigraphy and sequence stratigraphy of the sandy formations have been the subject of considerable study in recent years (engkilde & surlyk 1993; engkilde 1994; dam & surlyk 1995; engkilde & surlyk 2003, this volume; koppelhus & dam 2003, this volume), the precise age, the nature of the boundaries and the depositional history of the sortehat formation remain poorly understood. the aim of this paper, therefore, is to present the results of a detailed palynological study of these strata that was undertaken in close co-operation with a sedimentological and sequence stratigraphic study (hansen 1999). in particular, this paper focuses on establishing the age of the sortehat formation, elucidation of the stratigraphic significance of the formation boundaries and contributing to an understanding of the depositional history of the sortehat formation. regional setting and stratigraphy the middle jurassic sortehat formation is exposed in the jameson land basin in the southernmost part of the north–south-trending, failed-rift system in east greenland (fig. 1). rifting was initiated in the late palaeozoic, and 778 25 km 100 km illoqqortoormiut 22°w24°w 72¡ 72°n 71°n sortehat formation study locality scoresby sund jameson land li ve rp oo l l an d scoresby land kong oscar fjord pelion lepidopteriselv enhjørningen dal albuen vardekløft primulaelv harris fjeld h u rr y in le t sortehat dusén bjerg fle m in g fjo rd car lsb er g f jor d liaselv i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i ii i i i i i i i i i i i i i i hold with hope geographical society ø gauss halvø traill ø kejser franz joseph fjord kong oscar fjord scoresby sund jameson land milne land liverpool land 72°n 71°n 73°n 24°wa b n g re en la nd 25 km normal fault inferred cross-fault i i g re en la nd jameson land basin fig. 1. a: map showing the outcrop of the sortehat formation in the jameson land basin; the outcrop pattern is aligned roughly north–south, parallel to the basin axis. b: structural outline of the southern part of the east greenland rift basin. the jameson land basin was bounded by faults to the east, west and north, and by the liverpool land high towards the east. the nature of the southern boundary is unknown. the structural setting suggests an elongate semi-enclosed basin during deposition of the sortehat formation. modified from engkilde & surlyk (1993). from late permian through mesozoic times the basin was characterised by relatively uniform thermal subsidence interrupted by periods of faulting (surlyk et al. 1981; surlyk 1990a). the basin was bounded to the west by a major, approximately north–south-trending fault zone and to the east by faults and the elongated nne–ssw-trending liverpool land high (fig. 1; surlyk et al. 1981; surlyk 1990a). to the north, the basin was bounded by a number of nw–se-trending cross-faults in kong oscar fjord. the southern boundary is unknown but the basin probably extended further south under the present-day scoresby sund. the original extent of the sortehat formation is not known in detail, but during middle jurassic time it probably covered the present land area of jameson land and scoresby land (fig. 1). it was deposited in an elongate, semi-enclosed seaway, connected in the south with the epeiric jurassic seaway between greenland and norway (surlyk et al. 1981). the sortehat formation, as adopted here, was originally erected as the lower member (the sortehat member) of the vardekløft formation (surlyk et al. 1973). it was raised to the status of formation by surlyk (1990b, fig. 3); formal definition of the sortehat formation as the uppermost formation of the neill klinter group was undertaken by dam & surlyk (1998). the type locality of the sortehat formation at sortehat (fig. 1a) is identical to that of the former sortehat member (surlyk et al. 1973). 779 fig. 2. lithostratigraphy of the jurassic of jameson land, east greenland showing the stratigraphic position of the sortehat formation in the uppermost neill klinter group. modified from surlyk (2003, this volume, fig. 5). the boundary between the ostreaelv formation and the sortehat formation is sharp and erosive in the northern proximal part of the basin and truncates shoreface deposits (indicated by the wavy line). at southern localities, the offshore mudstones of the sortehat formation gradationally overlie transgressive lower shoreface to offshore deposits. the sortehat formation thins towards the south. the heterolithic unit in the lowermost part of the formation also thins towards the south. the upper boundary of the sortehat formation is erosional (wavy line) and may record a minor hiatus. thickness (m) formation 100–400 60–110 290–420 rhaetian hettangian sinemurian pliensbachian toarcian lo w er m id dl e ju ra ss ic tr ia s. u pp er aalenian bajocian bathonian callovian oxfordian kimmeridgian volgian groupchronostratigraphy formation sortehat ostreaelv gule horn rævekløft kap stewart ja m es on l an d supergroup scoresby sund hall bredning vardekløft neill klinter raukelv hareelv olympen fossilbjerget pelion r hæ te lvprimulaelv innakajik sandstone sandstone/mudstone heterolith mudstone pelion fm ostreaelv fm sortehat fm ? ? ns the sortehat formation overlies the sandstone-dominated ostreaelv formation (figs 2, 3) which records deposition within a shallow wave, storm and tidally influenced marine embayment (dam & surlyk 1995, 1998). the boundary with the overlying black mudstones of the sortehat formation is distinct at all localities but changes from a sharp ravinement surface in northern localities to a gradational drowning surface in southern localities (figs 3, 4; hansen 1999). the sortehat formation is 60–100 m thick and is overlain by the sandy, marine pelion formation (fig. 2); the boundary is sharp throughout the basin. sedimentologically, the boundary between the ostreaelv and sortehat formations represents a landwards shift in facies and a rise in relative sea level. the contact is interpreted as a transgressive erosional surface (surlyk 1990a, b; hansen & surlyk 1994, hansen 1999) and marks a basinwide flooding event within an overall transgressive period. the upper boundary of the sortehat formation represents a seawards shift in facies related to a fall in sea level and is interpreted as a marine erosional surface formed during forced regression (surlyk 1990a, b; engkilde & surlyk 1993). facies and depositional setting the sortehat formation consists of dark grey to black mudstones with subordinate heterolithic levels (figs 3, 4). the formation thins southwards away from the northern basin margin, from 100 m at enhjørningen dal to 60 m at albuen (fig. 4). a discrete, heterolithic unit of interbedded mudstones, sandstones and siltstones, 20 m thick, occurs at the base of the formation at the northern locality of enhjørningen dal; this unit thins to 8 m at the type locality of the sortehat formation and ultimately wedges out towards the south (figs 3, 4). mudstone facies the mudstones of the sortehat formation are darker and more fissile in the lowermost part of the formation and generally become lighter upwards, changing from black at the base to dark grey at the top. the dark grey mudstones, which dominate the formation, commonly appear structureless but locally show bioturbated fabrics, including subhorizontal traces such as curvolithus isp. and planolites-like burrows and, near the top of the formation, some vertical traces referred to diplocraterion isp. the black mudstones, in contrast, show a welldeveloped coarse parallel lamination formed by an alternation of thin sand/siltstone layers and mudstone layers. the mudstones contain plant debris and calcareous concretions are present locally. macrofossils include belemnites and ostreid bivalves. both mudstone facies are broadly interpreted as having been deposited from suspension below wave base in an offshore environment. the lack of bioturbation in the black mudstone facies suggests that the sea floor was periodically inhospitable, probably due to poor oxygenation. sandstone facies the interbedded sandstone layers from the heterolithic levels are very fineto medium-grained; they locally show well-developed wave ripples but more commonly appear as lenses and streaks of sandstone, 2–15 mm thick, reflecting incipient ripple development (fig. 3c; cf. the ‘incipient lenses’ of de raaf et al. 1977). laterally persistent sandstone layers, 5–30 cm thick, are present locally and show hummocky cross-stratification (fig. 3c). the trace fossils ophiomorpha nodusa and pelecypodichnus amygdaloides occur in some of the sandstone layers. the basal heterolithic unit records deposition in the upper offshore – offshore transition zone, influenced by storm sand deposition on a muddy shelf. the persistent hummocky cross-stratified sand sheets represent higher energy storm events above storm wave base. the bioturbation associated with the sandstones indicates wellaerated bottom-water conditions, at least on occasion. depositional setting the preliminary interpretation is that the black to dark grey mudstones that dominate the formation were deposited from suspension. in combination with the thin storm sand layers, the mudstones indicate deposition in an offshore environment, probably a muddy, shallow epeiric sea. the basal thick heterolithic unit at enhjørningen dal and sortehat records northwards shoreface retreat prior to final drowning of the entire basin. a detailed sedimentological and sequence stratigraphic analysis of the sortehat formation was presented by hansen (1999). 780 781 a fig. 3. a: the sortehat formation at the southernmost locality, albuen, neill klinter. the sortehat formation consists mainly of dark grey to black mudstones deposited in an offshore environment. the boundary with the underlying fully marine sandy ostreaelv formation (arrow) is a particularly distinctive facies boundary within the mesozoic succession of the east greenland rift basin and represents a marine flooding event. the profile shown is c. 80 m thick. b: the lower, heterolithic unit of the sortehat formation at the northern locality of enhjørningen dal (basal boundary marked by arrow; c. 15 m of the sortehat formation illustrated). c: in contrast to the shaly appearance at the southern locality of albuen (see fig. 3a), this 20 m thick heterolithic unit at enhjørningen dal consists of stacked coarsening-upwards units (arrow); the example illustrated here passes up from mudstone (m) to hummocky cross-stratified sandstone (hcs). these stacked units probably record repeated slowing of transgression and shoreline progradation at the northern basin margin. measuring rule (centre left) is 20 cm long. b c previous work the age of the sortehat formation is not well known. belemnites have been collected from the formation but to date have not been systematically identified. according to dam & surlyk (1998), belemnites from the upper levels of the underlying ostreaelv formation were collected by rosenkrantz (1934) and studied by doyle (1991); the belemnite ‘parabrachybelus’ subaduncatus from this level probably has a range restricted to the latest toarcian levesquei zone (doyle 1991). the ammonite cranocephalites borealis has been found in the basal beds of the pelion formation, overlying the sortehat formation, indicating an age not younger than early late boreal bajocian for these beds (surlyk et al. 1973; callomon 1993). the dark shales of the sortehat formation have been subjected to three palynological studies and a geochemical study (fensome 1979; lund & pedersen 1985; krabbe et al. 1994; underhill & partington 1994) although none of these studies were based on a comprehensive, closely-spaced sampling programme. three samples from the sortehat formation were analysed for dinoflagellate cysts by fensome (1979). one sample (144112) 782 9 8 7 6 0 sentusidinium pelionense assemblage zone nannoceratopsis gracilis – nannoceratopsis senex assemblage zone botryococcus sp. assemblage zone perinopollenites elatoides assemblage zone 0 100 m 90 80 70 60 50 40 30 20 10 60 m 50 40 30 20 10 60 m 50 40 30 20 10 0 mudstone pebbles belemnites bivalves ammonites logs concretions degree of bioturbation mudstone with sand lenses sandstone sandstone with clay laminae albuen sortehat enhjørningen dal s n < 75 km >< 40 km > cross-bedded sandstone hummocky cross-stratified sandstone pelion fm sortehat fm ostreaelv fm fig. 4. correlation between the three localities albuen, sortehat and enhjørningen dal showing the palynological zonation (assemblage zones 6–9). the upper part of the sortehat formation and the overlying pelion formation at the type locality (sortehat) have been removed by recent erosion. the base and top of the sortehat formation are indicated by the dotted line. was from the type section at sortehat, and two (144229 and 144231) were from a locality north of dusén bjerg. sample 144112 yielded a sparse assemblage of well-preserved palynomorphs, including nannoceratopsis gracilis. sample 144229 yielded a dinoflagellate cyst assemblage dominated by n. gracilis. the assemblage in sample 144231 was dominated by the acritarch veryhachium sortehatense. fensome (1979) concluded that the palynomorph assemblages determined from the three samples from the sortehat formation did not allow for accurate dating. it is worth noting that the dinoflagellate cyst sentusidinium pelionense was not found in any of the three samples from the sortehat formation, but was common in one sample (144111) from the overlying pelion formation. lund & pedersen (1985) presented the results of a palynological study concerning the neill klinter and vardekløft groups and the lower part of the hareelv formation. four samples (142832–35) are from the sortehat formation. these samples yielded abundant sentusidinium pelionense but nannoceratopsis gracilis was not found, whereas the pollen perinopollenites elatoides was abundant. sample 142833 had the lowest number of marine cysts, whereas botryococcus was common in 142832 and 142835. on the basis of these data, lund & pedersen (1985) suggested a middle–late bajocian age for the sortehat formation. the neill klinter and vardekløft groups were also studied palynologically by underhill & partington (1994). the material on which their study was based was sampled at liaselv (their section 1), vardekløft (section 2), and harris fjeld/primulaelv at neill klinter, the western slope of hurry inlet (section 3; fig. 1). twenty samples from the sortehat formation were analysed from their sections 1 and 2 (7 samples from section 1, 13 samples from section 2; their fig. 10). an aalenian–bajocian age was proposed for the sortehat formation. on the basis of these data, underhill & partington (1994) suggested that the aalenian–bajocian record was essentially complete, without apparent biostratigraphic or sedimentological evidence of the ‘mid-cimmerian event’ known from the north sea area. underhill & partington (1994) concluded that the boundary between the sortehat formation and the underlying ostreaelv formation does not represent an important uncomformity. this is confirmed by recent work (koppelhus & dam 2003, this volume) on the uppermost part of the ostreaelv formation which has been referred to the late toarcian – early aalenian on the basis of the palynological assemblages. the organic geochemistry and the palynofacies of the sortehat formation were discussed by krabbe et al. (1994). based on the palynofacies study, the succession was divided into three facies: (1) botryococcus-dominated, (2) spore/pollen and brown/blackwood and (3) blackwood-dominated, few spore/pollen. these results, together with the geochemical data, suggest an increase in the salinity of the depositional environment with time (krabbe et al. 1994). materials and methods this study is primarily based on material collected by the authors in the 1993 and 1994 field seasons. the sortehat formation was investigated along a north–southtrending profile through the jameson land basin, parallel to the basin axis. sections were sampled and studied at four localities: albuen (neill klinter along hurry inlet) and sortehat (the type locality of the formation) in the south and enhjørningen dal and pelion in the northern part of jameson land (fig. 1; appendix 1). the boundary between the ostreaelv and the sortehat formations is a well-defined stratigraphic surface and was used as a datum for the sections measured at outcrop; the structural dip of the succession is negligible. the outcrop sections were measured by jacob staff in metres relative to this surface. altimeter readings for the datum surface are listed in appendix 1; sample locations are thus referred to height above sea level, being the sum of the datum altitude (measured by altimeter) and the measured section thickness above the datum. samples taken from core from the borehole at the type locality of the sortehat formation are related to an arbitrary datum (base of cored section) within the upper ostreaelv formation (appendix 1). approximately 300 samples were processed at the palynological laboratory of the former geological survey of greenland using standard techniques (nøhr-hansen 1993). the palynomorphs were studied using a transmitted light microscope. for each sample, 200 specimens were counted, and all species were registered in the range chart program sis. the palynomorphs illustrated in plates 1–6 are from the borehole at the type locality of the sortehat formation and from a section at lepidopteriselv; the latter has not been used in the correlation between the three other localities. all the palynomorph taxa recorded in the samples are listed in appendix 2 with author attributions and dates, and the slides are stored in the collections of the geological survey of denmark and greenland. 783 palynological zones: definition the data presented here form part of a broader study encompassing the entire neill klinter group. nine palynological assemblage zones (1–9) have been recognised in the group; assemblage zones 1–6 from the rævekløft, gulehorn and ostreaelv formations are presented in a companion paper (koppelhus & dam 2003, this volume). the three assemblage zones of the sortehat formation, based on the occurrence of miospores, dinoflagellate cysts and freshwater algae, are named from below: (7) the botryococcus assemblage zone, (8) the nannoceratopsis gracilis – nannoceratopsis senex assemblage zone and (9) the sentusidinium pelionense assemblage zone. assemblage zone 7 is also recognised in the upper levels of the ostreaelv formation and is described briefly in koppelhus & dam (2003, this volume), but is defined herein. assemblage zone 7: botryococcus new assemblage zone occurrence. albuen 438.5–443.5 m enhjørningen dal 425.35–445 m pelion 550–567 m sortehat (core) 27.82–36.36 m in the cored section from sortehat, this assemblage occurs in the lower levels of the sortehat formation, the base being immediately above the lower boundary of the formation (figs 4, 5). at albuen, the assemblage is represented in the uppermost few metres of the ostreaelv formation and extends up into the sortehat formation (figs 4, 6; fig. 6 faces page 794). at enhjørningen dal, assemblage zone 7 is restricted to the lower sortehat formation although here the base is some 6 m above the lower boundary of the sortehat formation (figs 4, 7; fig. 7 faces page 795). assemblage zone 7 occurs within the lower levels of sequence sq7 of dam & surlyk (1995, 1998). reference section. sortehat (core), 27.82 m (sample 303143-73) – 36.36 m (sample 303143-62; figs 4, 5). additional sections. albuen, 438.5 m (sample 397452) – 443.5 m (sample 397468; figs 4, 6). enhjørningen dal, 425.35 m (sample 398341) – 445 m (sample 398417; figs 4, 7). base. the base of the assemblage is placed at the first sample in which botryococcus spp. overwhelmingly dominates the assemblage and the dinoflagellate cysts nannoceratopsis gracilis and nannoceratopsis senex and dinoflagellate cysts in general become rare. in the albuen section (fig. 6), this event coincides with the first co-occurrence of callialasporites dampieri (pollen) and mendicodinium groenlandicum (dinoflagellate cyst) although in other sections (e.g. sortehat, enhjørningen dal; figs 5, 7) these species first occur together some metres below the botryococcus spp. influx. top. the upper boundary is defined by the uppermost sample showing the botryococcus-dominated assemblage. above this level, botryococcus spp. are scarce and nannoceratopsis gracilis and n. senex become abundant once more. characteristics. the assemblage is characterised by the overwhelming dominance of the freshwater alga botryococcus spp. and the scarcity of dinoflagellates. pollen species such as perinopollenites elatoides, cerebropollenites macroverrucosus and bisaccate pollen are also abundant. suggested age. aalenian (see discussion below). palaeoenvironment. the palynomorph assemblage indicates a brackish marine environment. the abundant botryococcus and the common spores and pollen reflect a significant allochthonous terrestrial input related to transgression (see later discussion). remarks. the botryococcus assemblage is, as the name indicates, dominated by botryococcus spp. (plate 4, fig. 4), but spores and pollen also play an important role. the spore flora is diverse, but there are only few specimens of each species, whereas pollen species are less diverse but occur abundantly, such as perinopollenites elatoides (plate 2, fig. 5), bisaccate pollen, cerebropollenites macroverrucosus (plate 3, fig. 3) and corollina torosus (plate 3, fig. 2). some of the less abundant pollen are stratigraphically significant, such as quadraeculina anellaeformis (plate 3, fig. 5) and callialasporites dampieri (plate 2, fig. 1). acritarchs and dinoflagellate cysts occur only rarely. among the latter are dissilodinium sp. (plate 6, fig. 4), mendicodinium sp., pareodinia halosa (plate 4, fig. 6), mancodinium semitabulatum (plate 6, fig. 1) and a few specimens of nannoceratopsis senex (plate 5, fig. 2) and n. gracilis (plate 5, fig. 1). 784 785 so rt eh at middle jurassic lower jurassic aalenian toarcian ostreaelv formationsortehat formation 77 .9 8 73 .1 3 69 .5 0 65 .2 4 61 .3 1 57 .4 1 53 .8 6 49 .4 8 45 .0 9 40 .1 1 36 .3 6 32 .0 6 29 .5 6 27 .8 2 26 .2 8 12 .6 5 30 31 43 -1 1 30 31 43 -1 6 30 31 43 -2 0 30 31 43 -2 4 30 31 43 -2 8 30 31 43 -3 3 30 31 43 -3 7 30 31 43 -4 3 30 31 45 -4 9 30 31 43 -5 5 30 31 43 -6 2 30 31 43 -6 7 30 31 43 -6 9 30 31 43 -7 3 30 31 43 -7 5 30 31 43 -7 7 1baculatisporites spp. 2manumia delcourtii 3tripartina variabilis 4staplinisporites caminus 5retitriletes semimuris 6retitriletes clavatoides 7deltoidospora spp. 8retitriletes spp. 9leptolepidites bossus 10striatella jurassica 11uvaesporites puzzlei 12taurocusporites verrucatus 13concavissimisporites spp. 14sestrosporites pseudoalveolatus 15ischyosporites spp. 16leptolepidites spp. 17ischyosporites variegatus 18striatella parva 19lycopodiacidites rugulatus 20densoisporites scanicus 21foraminisporis jurassicus 22densoisporites velatus 23polycingulatisporit triangularis 24megaspore spp. 25neoraistrickia spp. 26striatella spp. 27retitriletes austroclavatoides 28kraeuselisporites reissingeri 29apiculatisporites spp. 30densosporites spp. 31chasmatosporites hians 32spheripollenites subgranulatus 33bisaccate spp. 34perinopollenites elatoides 35pinuspollenites minimus 36cerebropollenites macroverrucosus 37corollina torosus 38vesicaspora fuscus 39quadraeculina anellaeformis 40cerebropollenties thiergartii 41chasmatosporites major 42callialasporites dampieri 43callialasporites turbatus 44exesipollenites tumulus 45callialasporites trilobatus 46araucariacites australis 47chasmatosporites apertus 48eucommiidites troedsonii 49monosaccate spp. 50dodekovia tegilla 51wallodinium laganum 52nannoceratopsis gracilis 53scriniocassis weberii 54nannoceratopsis ambonis 55nannoceratopsis senex 56mancodinium semitabulatum 57dinocyst spp. 58mendicodinium groenlandicum 59dissiliodinium spp. 60pareodinia halosa 61mendicodinium reticulatum 62mendicodinium spp. 63kallosphaeridium spp. 64pareodinia ceratophora 65sentusidinium pelionense 66nannoceratopsis plegas 67susadinium scrofoides 68phallocysta eumekes 69andreedinium arcticum 70phallocysta elongata 71scriniocassis spp. 72acritarch spp. 73limbicysta bjaerkei 74veryhachium formosum 75veryhachium sortehatense 76lecaniella foveata 77leiosphaeridia spp. 78botryococcus spp. 79tasmanites spp. 80foraminifera spp. ? r ? ? ? a lp ha be tic al s pe ci es li st 72 a cr ita rc h sp p. 69 a n d re ed in iu m a rc ti cu m 29 a p ic u la ti sp or it es s pp . 46 a ra u ca ri a ci te s a u st ra lis 1 b a cu la ti sp or it es s pp . 33 bi sa cc at e sp p. 78 b ot ry oc oc cu s sp p. 42 c a lli a la sp or it es d a m p ie ri 45 c a lli a la sp or it es t ri lo b a tu s 43 c a lli a la sp or it es t u rb a tu s 36 c er eb ro p ol le n it es m a cr ov er ru co su s 40 c er eb ro p ol le n ti es t h ie rg a rt ii 47 c h a sm a to sp or it es a p er tu s 31 c h a sm a to sp or it es h ia n s 41 c h a sm a to sp or it es m a jo r 13 c on ca vi ss im is p or it es s pp . 37 c or ol lin a t or os u s 7 d el to id os p or a s pp . 20 d en so is p or it es s ca n ic u s 22 d en so is p or it es v el a tu s 30 d en so sp or it es s pp . 57 d in oc ys t sp p. 59 d is si lio d in iu m s pp . 50 d od ek ov ia t eg ill a 48 e u co m m iid it es t ro ed so n ii 44 e xe si p ol le n it es t u m u lu s 80 fo ra m in ife ra s pp . 21 fo ra m in is p or is j u ra ss ic u s 15 is ch yo sp or it es s pp . 17 is ch yo sp or it es v a ri eg a tu s 63 k a llo sp h a er id iu m s pp . 28 k ra eu se lis p or it es r ei ss in ge ri 76 le ca n ie lla f ov ea ta 77 le io sp h a er id ia s pp . 9 le p to le p id it es b os su s 16 le p to le p id it es s pp . 73 li m b ic ys ta b ja er ke i 19 ly co p od ia ci d it es r u gu la tu s 56 m a n co d in iu m s em it a b u la tu m 2 m a n u m ia d el co u rt ii 24 m eg as po re s pp . 58 m en d ic od in iu m g ro en la n d ic u m 61 m en d ic od in iu m r et ic u la tu m 62 m en d ic od in iu m s pp . 49 m on os ac ca te s pp . 54 n a n n oc er a to p si s a m b on is 52 n a n n oc er a to p si s gr a ci lis 66 n a n n oc er a to p si s p le ga s 55 n a n n oc er a to p si s se n ex 25 n eo ra is tr ic k ia s pp . 64 pa re od in ia c er a to p h or a 60 pa re od in ia h a lo sa 34 pe ri n op ol le n it es e la to id es 70 ph a llo cy st a e lo n ga ta 68 ph a llo cy st a e u m ek es 35 pi n u sp ol le n it es m in im u s 23 po ly ci n gu la ti sp or it t ri a n gu la ri s 39 q u a d ra ec u lin a a n el la ef or m is 27 r et it ri le te s a u st ro cl a va to id es 6 r et it ri le te s cl a va to id es 5 r et it ri le te s se m im u ri s 8 r et it ri le te s sp p. 71 s cr in io ca ss is s pp . 53 s cr in io ca ss is w eb er ii 65 s en tu si d in iu m p el io n en se 14 s es tr os p or it es p se u d oa lv eo la tu s 32 s p h er ip ol le n it es s u b gr a n u la tu s 4 s ta p lin is p or it es c a m in u s 10 s tr ia te lla j u ra ss ic a 18 s tr ia te lla p a rv a 26 s tr ia te lla s pp . 67 s u sa d in iu m s cr of oi d es 79 ta sm a n it es s pp . 12 ta u ro cu sp or it es v er ru ca tu s 3 tr ip a rt in a v a ri a b ili s 11 u va es p or it es p u z z le i 74 ve ry h a ch iu m f or m os u m 75 ve ry h a ch iu m s or te h a te n se 38 ve si ca sp or a f u sc u s 51 w a llo d in iu m l a ga n u m 75 50 25 system stage palynological assembl. zones lithostratigraphy (m) sample height sample number 9 8 7 6 u nc er ta in de te rm in at io n ve ry r ar e r ar e fe w c om m on a bu nd an t ? r fi g. 5 . c h ar t sh o w in g th e d is tr ib u tio n o f p al yn o m o rp h s fr o m t h e co re d b o re h o le a t so rt eh at , th e ty p e lo ca lit y o f th e so rt eh at f o rm at io n . assemblage zone 8: nannoceratopsis gracilis – nannoceratopsis senex new assemblage zone occurrence. albuen 444.5–460 m enhjørningen dal 446–500 m pelion 573–577 m sortehat (core) 40.11–53.86 m assemblage zone 8 occurs within the sortehat formation; it thus falls within sequence sq7 of dam & surlyk (1995, 1998). reference section. sortehat (core), 40.11 m (sample 303143-55) – 53.86 m (sample 303143-37; figs 4, 5). additional sections. albuen, 444.5 m (sample 397469) – 460 m (sample 397474; figs 4, 6). enhjørningen dal, 446 m (sample 395625) – 500 m (sample 398442; figs 4, 7). base. the lower boundary is placed at the first sample in which botryococcus spp. is rare and nannoceratopsis gracilis and n. senex are abundant. top. this is defined by the last sample showing the characteristic assemblage (see below), above which level nannoceratopsis gracilis becomes less common and sentusidinium pelionense is the most common dinoflagellate cyst. characteristics. botryococcus spp. is rare in this assemblage whereas nannoceratopsis gracilis is abundant and there is a general increase in diversity and abundance of dinoflagellate cysts relative to the underlying zone. suggested age. aalenian (see discussion below). palaeoenvironment. the palynological data indicate a brackish marine environment. remarks. dinoflagellate cysts are abundant, whereas botryococcus spp. becomes rare at 40.11 m in the type section at sortehat (fig. 5). there are a few acritarchs, such as veryhachium sorthatense (plate 4, fig. 2), and some of the same spores and pollen as observed in the botryococcus assemblage zone (plates 1, 2). assemblage zone 9: sentusidinium pelionense new assemblage zone occurrence. albuen 465–502 m enhjørningen dal 502–516.75 m sortehat (core) 57.41–77.98 m assemblage zone 9 occurs in the upper sortehat formation in the sortehat section; note that the sampled interval in this section is restricted to the sortehat formation (figs 4, 5). at albuen and enhjørningen dal, the assemblage extends through the upper sortehat formation and persists up into the overlying pelion formation (figs 4, 6, 7). assemblage zone 9 is thus characteristic of the upper levels of sequence sq7 of dam & surlyk (1995, 1998), and at albuen and enhjørningen dal spans the sequence boundary at the base of the pelion formation and extends into sequence p1 of engkilde & surlyk (2003, this volume). reference section. sortehat (core), 57.41 m (sample 303143-33) – 77.98 m (sample 303143-11; figs 4, 5). additional sections. albuen, 465 m (sample 397475) – 502 m (sample 397498; figs 4, 6). enhjørningen dal, 502 m (sample 395679) – 516.75 m (sample 398448; figs 4, 7). base. the base of the zone is placed at the sample in which sentusidinium pelionense is the most common dinoflagellate cyst; nannoceratopsis gracilis and n. senex are absent or rare. top. the upper boundary of the assemblage zone is not defined here. the sentusidinium pelionense assemblage zone extends from the upper part of the sortehat formation into the lowermost beds of the pelion formation. the full extent of the zone within the pelion formation is not known. it is likely, however, that the top of the zone occurs within the lower levels of the pelion formation (s. piasecki, personal communication 1997). characteristics. sentusidinium pelionense is abundant. in some of the investigated sections, s. pelionense is accompanied by nannoceratopsis gracilis in samples at the boundary between the two palynomorph assemblage zones. in other sections, they do not overlap, i.e. n. gracilis is replaced by s. pelionense up-section. 786 suggested age. aalenian – ?early bajocian (see discussion below). palaeoenvironment. the palynological data indicate that most of the organic material came from a brackish marine source. remarks. the assemblage contains fewer spores and pollen than the underlying assemblage zone. the most abundant dinoflagellate cysts are sentusidinium pelionense (plate 6, fig. 5) and pareodinia halosa (plate 4, fig. 6) although a few specimens of phallocysta eumekes were recorded. the acritarch limbicysta bjaerkei was also found. palynological results sortehat at the type locality, the sortehat formation is a minimum of 50 m thick; the top of the formation is not seen due to recent erosion. seventy samples were collected and seventeen of these are shown on the distribution chart (fig. 5). the palynological assemblages from the entire section fall into three distinct assemblages. there is a distinct change from the perinopollenites elatoides zone (assemblage zone 6, described by koppelhus & dam 2003, this volume) in the underlying ostreaelv formation to the botryococcus assemblage zone (assemblage zone 7), which is overwhelmingly dominated by the freshwater alga botryococcus. this change is first observed at 27.82 m (sample 303143-73) just above the lithological boundary between the sortehat and ostreaelv formations at 27.6 m. the perinopollenites elatoides zone is characterised by the first appearance of staplinisporites caminus, sestrosporites pseudoalveolatus, phallocysta eumekes and wallodinium laganum; callialasporites dampieri makes its first appearance near the top of the zone. the data from the three ostreaelv formation samples from the sortehat section do not show exactly the same pattern although s. caminus and c. dampieri have their first appearance in the uppermost sample from the ostreaelv formation and the dinoflagellate cyst w. laganum is very common in sample 303143-80 at 6.64 m (not shown on fig. 5), approximately 20 m below the boundary between the ostreaelv and sortehat formations. in sample 303143-73 at 27.82 m, 0.22 m above the base of the sortehat formation, the freshwater alga botryococcus appears and the acme extends for approximately 9 m before botryococcus disappears and the dinoflagellate cyst nannoceratopsis gracilis begins to dominate the assemblage. in three samples (303143-75, 303143-73 and 303143-69) an unidentified dinocyst (dinocyst sp.) is common to abundant; this dinoflagellate cyst is similar to one described as dinoflagellate sp. indet. 2 from callovian deposits from spitsbergen where it is said to occur in profusion in some assemblages (bjærke 1980). this species was also common in the lowermost sample from the pelion locality (fig. 8). n. gracilis continues to be the dominant dinoflagellate for 14 m to 53.86 m (sample 303143-37) above which there is an acme of sentusidinium pelionense together with abundant pareodinia halosa. albuen the sortehat formation at albuen covers about 60 m and 41 samples have been investigated from this interval (fig. 6). the boundary between the ostreaelv and sortehat formations is at 440 m, but the palynomorph assemblages change between sample 405449 at 434 m and sample 397452 at 438.6 m, several metres below the top of the ostreaelv formation, where botryococcus becomes common, dinoflagellate cysts become rare (mendicodinium groenlandicum, nannoceratopsis gracilis and nannoceratopsis senex are present but rare), and pollen are more common than spores. at 443.5 m in sample 397468, 3.5 m above the base of the sortehat formation, botryococcus is abundant for the last time; above this level n. gracilis and n. senex become abundant and are accompanied by sentusidinium pelionense. s. pelionense is only common in one sample (397471), and pareodinia halosa is common in two other samples, 397475 and 347477. this assemblage changes between 460 m and 465 m, above which level s. pelionense is the only common dinoflagellate; this species disappears above 496 m, within the lower levels of the pelion formation. the overall dinoflagellate cyst diversity decreases from the upper part of the nannoceratopsis gracilis – nannoceratopsis senex assemblage zone through the sentusidinium pelionense assemblage zone in the upper part of the sortehat formation. a few dinoflagellate cysts have been found in most of the samples, and pollen are more common than spores. the samples that span the boundary between the sortehat formation and the overlying pelion formation yielded an impoverished palynomorph assemblage without any age-diagnostic species. 787 enhjørningen dal enhjørningen dal is a composite section spanning the top of the ostreaelv formation, the entire sortehat formation and the lowermost part of the pelion formation (fig. 7). the boundary between the ostreaelv and sortehat formations is at 418 m; the first change in the palynomorph assemblages is observed in sample 398341 at 425.35 m. this lower interval (belonging to the uppermost part of the ostreaelv formation and the lowermost part of the sortehat formation) is completely dominated by the dinoflagellate species nannoceratopsis gracilis and nannoceratopsis senex, and by the pollen perinopollenites elatoides. this interval is accordingly referred to the perinopollenites elatoides assemblage zone (assemblage zone 6; koppelhus & dam 2003, this volume) which is of late toarcian – early aalenian age. at 425.35 m (sample 398341), botryococcus appears in abundance and persists in large numbers up to sample 398417 at 445 m. from sample 395625 at 446 m to sample 398442 at 500 m, the palynomorph assemblage contains few botryococcus; n. gracilis and n. senex are very common in the lower part of this interval. in the upper part of the section, however, from sample 395657 at 476 m to the uppermost sample (398448) at 516.75 m, n. gracilis and n. senex are absent whereas other dinoflagellates are present but rare (e.g. mancodinium semitabulatum, phallocysta eumekes, susadinium scrofoides, parvocysta barbata, dissilodinium sp. and mendicodinium sp.). sentusidinium pelionense is abundant in samples 395679 and 395671, pareodinia halosa is common in sample 395671 and kallosphaeridium sp. is abundant in the uppermost sample (398448). among the pollen at this level, bisaccates and cerebropollenites macroverrucosus are especially abundant. trilete spores seem to decrease in diversity and density compared to the lower part of the section. pelion ten samples were investigated from the pelion locality (fig. 8). they cover the lowermost 27 m of the sortehat formation, which is approximately 110 m thick at this locality. generally, the palynomorph assemblages are rich but the preservation is poor. the assemblage from the lowermost 17 m of the formation (sample 339702 at 550 m to sample 339709 at 567 m) is dominated by botryococcus. in sample 339709, botryococcus has its last abundant appearance. in this interval, dinoflagellate cysts are present and locally common (nannoceratopsis gracilis, mendicodinium groenlandicum and mendicodinium sp.). the two uppermost samples (sample 339710 at 573 m and sample 339711 at 577 m) yielded an assemblage rich in dinoflagellate cysts including nannoceratopsis gracilis, nannoceratopsis triceras, nannoceratopsis plegas, nannoceratopsis triangulata and mancodinium semitabulatus and some acritarchs. spores and pollen are common to abundant throughout the 27 m section. discussion the palynological study reported here has implications both for the age and regional correlation of the sortehat formation and for the environmental understanding of the formation, in association with the detailed sedimentological and sequence stratigraphic study (hansen 1999). age and correlation on the basis of macrofauna in the underlying and overlying formations (see previous discussion), the age of the sortehat formation is constrained between the latest toarcian and the mid-bajocian; previous palynological work has indicated an aalenian–bajocian (underhill & partington 1994) or a middle–late bajocian age (lund & pedersen 1985). with a view to a more well-founded understanding of the stratigraphy of the sortehat formation, the results of this study are compared with published palynological data from the north atlantic region. although palynological assemblages from the toarcian–bajocian interval have been widely reported from this region, few of the sections have been independently dated using ammonites and the aalenian assemblages, in particular, are commonly dominated by terrestrial material. the aalenian stage is named after aalen in germany, where the lowest part of the ‘braunjura’ crops out at the northern edge of the swabian alps. although no palynological papers have been published on material from the type locality, a dinoflagellate cyst assemblage has been described from two cores from hausen in south-western germany, 40 km from the type locality, and the lithological units have been correlated with the ammonite zonation from the eastern swabian alb (feistburkhardt 1990). the palynological assemblages are rich in spores and pollen, and only 5–20% of the total 788 789 pe lio n middle jurassic aalenian sortehat formation 57 7. 00 57 3. 00 56 7. 00 56 5. 00 56 3. 00 55 9. 00 55 7. 00 55 6. 00 55 5. 00 55 0. 00 33 97 11 33 97 10 33 97 09 33 97 08 33 97 07 33 97 06 33 97 05 33 97 04 33 97 03 33 97 02 1kekryphalospora distincta 2striatella spp. 3retitriletes spp. 4deltoidospora spp. 5baculatisporites spp. 6retitriletes clavatoides 7retitriletes semimuris 8leptolepidites spp. 9ischyosporites variegatus 10sestrosporites pseudoalveolatus 11striatella parva 12manumia delcourtii 13staplinisporites caminus 14rogalskaisporites cicatricosus 15todisporites major 16stereisporites stereoides 17retitriletes austroclavatoides 18neoraistrickia spp. 19polycingulatisporit triangularis 20striatella seebergensis 21densoisporites scanicus 22tripartina variabilis 23striatella jurassica 24leptolepidites bossus 25taurocusporites segmentatus 26uvaesporites spp. 27spheripollenites subgranulatus 28callialasporites spp. 29callialasporites turbatus 30cerebropollenites macroverrucosus 31bisaccate spp. 32perinopollenites elatoides 33pinuspollenites minimus 34corollina torosus 35quadraeculina anellaeformis 36callialasporites minus 37vitreisporites pallidus 38callialasporites dampieri 39spheripollenites psilatus 40exesipollenites tumulus 41chasmatosporites major 42alisporites robusta 43chasmatosporites apertus 44striate spp. 45chasmatosporites hians 46monosulcites spp. 47cerebropollenties thiergartii 48callialasporites trilobatus 49callialasporites segmentatus 50araucariacites australis 51phallocysta spp. 52nannoceratopsis spp. 53dinocyst spp. 54nannoceratopsis gracilis 55mendicodinium groenlandicum 56nannoceratopsis senex 57mendicodinium spp. 58nannoceratopsis plegas 59nannoceratopsis triceras 60mancodinium semitabulatum 61nannoceratopsis triangulata 62leiosphaeridia spp. 63acritarch spp. 64botryococcus spp. 65miscellaneous spp. 66fungal spp. a lp ha be tic al s pe ci es li st 63 a cr ita rc h sp p. 42 a lis p or it es r ob u st a 50 a ra u ca ri a ci te s a u st ra lis 5 b a cu la ti sp or it es s pp . 31 bi sa cc at e sp p. 64 b ot ry oc oc cu s sp p. 38 c a lli a la sp or it es d a m p ie ri 36 c a lli a la sp or it es m in u s 49 c a lli a la sp or it es s eg m en ta tu s 28 c a lli a la sp or it es s pp . 48 c a lli a la sp or it es t ri lo b a tu s 29 c a lli a la sp or it es t u rb a tu s 30 c er eb ro p ol le n it es m a cr ov er ru co su s 47 c er eb ro p ol le n ti es t h ie rg a rt ii 43 c h a sm a to sp or it es a p er tu s 45 c h a sm a to sp or it es h ia n s 41 c h a sm a to sp or it es m a jo r 34 c or ol lin a t or os u s 4 d el to id os p or a s pp . 21 d en so is p or it es s ca n ic u s 53 d in oc ys t sp p. 40 e xe si p ol le n it es t u m u lu s 66 fu ng al s pp . 9 is ch yo sp or it es v a ri eg a tu s 1 k ek ry p h a lo sp or a d is ti n ct a 62 le io sp h a er id ia s pp . 24 le p to le p id it es b os su s 8 le p to le p id it es s pp . 65 m is ce lla ne ou s sp p. 60 m a n co d in iu m s em it a b u la tu m 12 m a n u m ia d el co u rt ii 55 m en d ic od in iu m g ro en la n d ic u m 57 m en d ic od in iu m s pp . 46 m on os u lc it es s pp . 54 n a n n oc er a to p si s gr a ci lis 58 n a n n oc er a to p si s p le ga s 56 n a n n oc er a to p si s se n ex 52 n a n n oc er a to p si s sp p. 61 n a n n oc er a to p si s tr ia n gu la ta 59 n a n n oc er a to p si s tr ic er a s 18 n eo ra is tr ic k ia s pp . 32 pe ri n op ol le n it es e la to id es 51 ph a llo cy st a s pp . 33 pi n u sp ol le n it es m in im u s 19 po ly ci n gu la ti sp or it t ri a n gu la ri s 35 q u a d ra ec u lin a a n el la ef or m is 17 r et it ri le te s a u st ro cl a va to id es 6 r et it ri le te s cl a va to id es 7 r et it ri le te s se m im u ri s 3 r et it ri le te s sp p. 14 r og a ls k a is p or it es c ic a tr ic os u s 10 s es tr os p or it es p se u d oa lv eo la tu s 39 s p h er ip ol le n it es p si la tu s 27 s p h er ip ol le n it es s u b gr a n u la tu s 13 s ta p lin is p or it es c a m in u s 16 s te re is p or it es s te re oi d es 44 s tr ia te s pp . 23 s tr ia te lla j u ra ss ic a 11 s tr ia te lla p a rv a 20 s tr ia te lla s ee b er ge n si s 2 s tr ia te lla s pp . 25 ta u ro cu sp or it es s eg m en ta tu s 15 to d is p or it es m a jo r 22 tr ip a rt in a v a ri a b ili s 26 u va es p or it es s pp . 37 v it re is p or it es p a lli d u s 57 5 56 5 55 5 system stage palynological. assembl. zones lithostratigraphy (m) sample height sample number 8 7 fi g. 8 . c h ar t sh o w in g th e d is tr ib u tio n o f p al yn o m o rp h s fr o m t h e lo ca lit y at p el io n . palynomorph content is microplankton. in the sortehat formation, where the assemblages are also rich in spores and pollen, microplankton form 25–40% of the total palynomorph content. in the material from southwestern germany, twelve of the eighteen dinoflagellate cysts recognised have also been identified in the material from the sortehat formation. however, none of the four stratigraphically significant species for the aalenian/bajocian boundary from hausen, carpathodinium sp., dissilodinium giganteum, dissilodinium sp. a, and durotrigia daveyi, have with certainty been found in the material from the sortehat formation. the closest succession to jameson land with known palynological data of aalenian age is the stø formation (unit c) from the møre basin, offshore mid-norway (smelror et al. 1994). ammonites have not been found to confirm the age, but a few foraminifera of little stratigraphic value have been used, together with the presence of the dinoflagellate cyst phallocysta eumekes, which is restricted to the latest early toarcian – aalenian in europe (riding & thomas 1992). spores, pollen and dinoflagellate cysts from unit c are similar to those from the botryococcus and nannoceratopsis gracilis – nannoceratopsis senex assemblage zones of the sortehat formation. two thin levels with botryococcus have been observed in the stø formation although not as rich in abundance as in the sortehat formation (smelror et al. 1994). moreover, the abrupt shift seen in the sortehat formation from an assemblage dominated by botryococcus to one dominated by the dinoflagellate cysts nannoceratopsis gracilis, nannoceratopsis senex and pareodinia halosa is not recognised in the møre basin. it should be noted, however that the botryococcus assemblage zone of the sortehat formation is not stratigraphically significant but reflects only the palaeoenvironmental conditions within the jameson land basin at the time of deposition (see discussion below). aalenian palynomorph assemblages have also been identified from arctic canada and south-east canada (johnson & hills 1973; davies 1983; bujak & williams 1977), barents sea (hammerfest basin, nordkap basin, and franz josef land; smelror & below 1992; smelror 1994), north-west scotland, north yorkshire and gloucestershire in england (riding 1983, 1984a, b, 1987; riding et al. 1991; riding & thomas 1992), sweden (guy-ohlson 1994; guy-ohlson & norling 1994), the danish subbasin (dybkjær 1991; seidenkrantz et al. 1993; poulsen 1994), øresund and the baltic sea (koppelhus & nielsen 1994; koppelhus & batten 1996), and north-west and south-west germany (prauss 1989). the first dinoflagellate cyst zonation of the jurassic of the canadian arctic was made by johnson & hills (1973). their nannoceratopsis gracilis range zone covers the toarcian–bajocian. they recorded the ammonite leioceras opalinum, but the only dinoflagellate cyst in common with the east greenland material is nannoceratopsis gracilis. bujak & williams (1977) erected the nannoceratopsis gracilis zone for the pliensbachian– aalenian and the mancodinium semitabulatum zone for the bajocian, from successions offshore south-eastern canada. several species are common to the assemblages from the sortehat formation and from the canadian arctic but the age has not been confirmed by any marine microor macrofauna. davies (1983) established eight zones covering the upper pliensbachian – callovian, also from the canadian arctic, of which the dapcodinium coalitum – phallocysta eumekes zone, of late toarcian – early bajocian age, contains species in common with the sortehat formation assemblages. davies (1983) stated that the macrofauna possibly indicates a toarcian – early bajocian age. in the barents sea area, seven dinoflagellate cyst zones have been recognised in the toarcian – lower oxfordian (smelror & below 1992). one of them, the dodekovia bulla – nannoceratopsis senex concurrent range-zone, has a number of species (nannoceratopsis gracilis, nannoceratopsis senex, nannoceratopsis triceras, scriniocassis weberi, susadinium scrofoides, pareodinia halosa and phallocysta eumekes) in common with the assemblages from the sortehat formation. the presence of the earliest aalenian opalinum zone on svalbard has been confirmed on the basis of ammonites. in the united kingdom, the aalenian palynomorph assemblages are divided into sub-biozone c (opalinum zone) and sub-biozone d (murchisonae and concavum zones) of the nannoceratopsis gracilis zone (riding & thomas 1992; the dsj10 and dsj11 zones of poulsen & riding 2003, this volume). these zones are based on the fad (first appearance datum) and the lad (last appearance datum) of specific dinoflagellate cyst species. of the stratigraphically significant species appearing in these zones, susadinium scrofoides, wallodinium laganum and other species of the genus parvocysta have also been found in the material from the sortehat formation. other common but stratigraphically less restricted species are nannoceratopsis gracilis, nannoceratopsis senex, mancodinium semitabulatum and scriniocassis weberi. the palynological zonation in the uk area has been related to the standard ammonitebased zones (riding & thomas 1992). 790 in skåne, southern sweden, the upper part of the rya formation and the lower part of the vilhelmsfält formation contain palynomorph assemblages dominated by terrestrial material deposited in a freshwater environment, although thin marine and brackish levels occur. palynological zone iii is suggested to be of aalenian age (guy-ohlson & norling 1994), and shows similarities to the palynomorph assemblages from the three zones of the sortehat formation. no ammonites have been found in the middle jurassic of southern sweden but a foraminifera zonation has been established. in the danish basin, rocks of toarcian and early aalenian age have been identified based on the occurrence of species of the genus parvocysta (poulsen 1994). a miospore and dinoflagellate cyst zonation was erected for the lower and middle jurassic of the danish basin (dybkjær 1991). the perinopollenites elatoides miospore zone of dybkjær (1991) covers sediments of aalenian – early bajocian age, whereas the equivalent dinoflagellate cyst zone includes nannoceratopsis gracilis and comprises sediments of late pliensbachian – bajocian age. similar palynomorph assemblages have been reported from the øresund area (koppelhus & batten 1996). a miospore zonation was recommended for the lower–middle jurassic of bornholm in the baltic area, where the callialasporites–perinopollenites zone from the bagå formation has many species in common with the sortehat formation (batten et al. 1994; koppelhus & nielsen 1994). dinoflagellate cysts (nannoceratopsis gracilis) have only been found in the lowermost part of this zone. the callialasporites–perinopollenites zone is defined by the first appearance of callialasporites and the dominance of perinopollenites elatoides. unfortunately, confirmatory ammonite or microfossil data is lacking in the danish area to support the proposed aalenian–bathonian age, which is based entirely on comparison to other palynological studies in europe. in the sentusidinium pelionense assemblage zone of the sortehat formation, spores and pollen are abundant, and dinoflagellate cysts are locally present. the dinoflagellate cyst sentusidinium pelionense has also been found in the jydegård formation on bornholm in the baltic sea (piasecki 1984; noe-nygaard et al. 1987). however, with the exception of the presence of s. pelionense, the sentusidinium pelionense assemblage zone of the sortehat formation is more closely comparable to assemblages from the norwegian and uk areas than to those of the danish basin and bornholm. palynomorph assemblages from the pliensbachian– callovian of north-west germany were described by prauss (1989). the palynomorph assemblage from the nannoceratopsis plegas zone of aalenian age, has a number of species in common with the three zones from the sortehat formation, such as nannoceratopsis gracilis, mancodinium semitabulatum, pareodinia halosa, dodekovia tegillata, scriniocassis weberi, phallocysta eumekes and kallospharidium sp. unfortunately the suggested aalenian age is not confirmed by ammonites. to conclude this review, the sortehat formation palynomorph assemblages clearly show close similarities to assemblages of inferred aalenian age in the north atlantic region. few of these studies, however, include independent ammonite data with which to accurately constrain the age of the strata. a notable exception, upon which the age assignment of the sortehat formation depends, is the work of riding (1982), woollam & riding (1983) and riding & thomas (1992) from the jurassic of the uk. according to the zonation presented by these workers, the stratigraphically important palynomorphs recorded from the sortehat formation are mancodinium semitabulatum, nannoceratopsis ambonis, phallocysta eumekes and nannoceratopsis plegas. m. semitabulatum, which was recorded throughout the sortehat formation, has a range of late pliensbachian – early bajocian (woollam & riding 1983; riding & thomas 1992). n. ambonis, which occurs in the lower half of the formation, was initially thought to range from the aalenian to the early bajocian (riding 1982; woollam & riding 1983) but its range has subsequently been extended to late pliensbachian – early bajocian (riding & thomas 1992). p. eumekes, which occurs throughout the sortehat formation, has a restricted range of late toarcian – aalenian in europe (riding & thomas 1992). however, the species was first described from bathonian strata in arctic canada where it has a range of toarcian–bathonian (dörhöfer & davies 1980) and is generally considered a more long-ranging species in boreal regions (riding 1984c). n. plegas is rare in the sortehat formation although its presence in the lower levels of the formation indicates an early aalenian age for this part of the succession (riding & thomas 1992). in association, therefore, these palynomorphs indicate an aalenian – ?early bajocian age for the sortehat formation. this study thus confirms the age determination of underhill & partington (1994), although it is clear from the above review that more detailed studies of both the miospore and the dinoflagellate cyst stratigraphies of the early middle jurassic are urgently needed to further resolve the stratigraphy of the sortehat formation. 791 a notable conclusion of this study is that there is no evidence, on palynological grounds, for a significant hiatus at, or near, the lower boundary of the sortehat formation (see also koppelhus & dam 2003, this volume), nor within the sortehat formation itself. these observations are particularly pertinent to regional studies of middle jurassic uplift in the north atlantic region (see underhill & partington 1994). the stratigraphic significance of the upper boundary of the sortehat formation is less well-constrained palynologically, although there is no direct evidence from this study of a major stratigraphic break at this surface. environmental implications of the palynological data the three palynomorph assemblage zones of the sortehat formation can be recognised and correlated throughout the jameson land basin (fig. 4). regionally, the botryococcus assemblage zone ranges in thickness from 9–17 m and spans the uppermost part of the ostreaelv formation and the lower part of the sortehat formation. the zone starts several metres below the boundary between the ostreaelv and sortehat formations in the southernmost locality at albuen, whereas the appearance of the zone coincides with the boundary at the sortehat type locality. farther north, at enhjørningen dal, this change in the palynomorph assemblage does not occur until 7 m above the formation boundary (figs 4, 7). at the northernmost locality of pelion, the change takes place 5 m above the boundary. the green alga botryococcus is a freshwater form (guy-ohlson 1992) with no stratigraphic value as it ranges from the carboniferous to the present (tyson 1995; batten & grenfell 1996). however, the abundance of botryococcus in the organic material from the lowermost sortehat formation in jameson land is notable. sedimentologically, the depositional shift from the ostreaelv formation to the sortehat formation marks a basinwide flooding event within an overall transgressive period. as the abundance of botryococcus is restricted to the section around the flooding surface (base of sortehat formation) and extends up to the maximum flooding surface, it seems to be linked to the flooding event (hansen 1999). the question is whether the abundance of botryococcus reflects in situ deposition under freshwater conditions that prevailed throughout the basin or if it represents an allochthonous accumulation of algae transported out to sea from inland freshwater environments. the continuous occurrence of dinocysts throughout the sortehat formation, albeit in reduced numbers in the botryococcus assemblage zone, suggests a persistent marine influence. indeed, the reduced number of marine dinoflagellate cysts in this zone may be in part an artefact related to the counting procedure such that the abundance of botryococcus and terrestrial sporomorphs tends to dilute the marine dinoflagellate cysts. given the co-occurrence of the freshwater alga botryococcus and the marine dinoflagellate cysts, the abundance of botryococcus is here interpreted as an allochthonous accumulation. seawards transport of material from inland areas suggests either a major freshwater input from rivers or inundation and erosion of coastal areas with seawards transport of landderived material during transgression. during marine flooding, the initial rise of base-level causes expansion of lakes on low-lying areas within the coastal plain (wells & coleman 1987; dominguez & wanless 1991; surlyk et al. 1995). as transgression proceeds, physical communication with the open sea is established, resulting in the possibility for a large influx of freshwater algae into the marine system. such an allochthonous origin related to transgression is consistent with the fact that the top of the botryococcus assemblage zone coincides with the maximum flooding surface and the end of the transgressive period (hansen 1999). a similar influx of botryococcus, related to a flooding event, has been described from the middle jurassic brent group of the north sea by williams (1992), who interpreted this as an allochthonous accumulation resulting from the flushing of a freshwater environment into a marine environment. the diversification of the dinoflagellate cysts in the basal part of the nannoceratopsis gracilis – nannoceratopsis senex assemblage zone is suggestive of a general increase in marine influence. this may be related to an increase in salinity during sea-level highstand following the transgression, as it is generally suggested that an increased diversity of dinoflagellate cysts points to more open marine conditions (gorin & steffen 1991; leckie et al. 1992). upwards within the nannoceratopsis gracilis – nannoceratopsis senex assemblage zone and through the sentusidinium pelionense assemblage zone, especially at albuen and partly at enhjørningen dal, there is a marked decrease in the diversity of dinoflagellate cyst species, although the assemblage still shows marine influence. a decrease in diversity indicates more stressed and unfavourable environmental conditions, often with unstable salinities (gorin & steffen 1991; leckie et al. 1992; tyson 1995). 792 although the strength of the marine signal partly increases up-section, the dinoflagellate cysts are characterised by forms tolerant of reduced salinities (e.g. nannoceratopsis gracilis, nannoceratopsis senex and sentusidinium pelionense; piasecki 1986; prauss & riegel 1989; krabbe et al. 1994). the palynological data therefore indicate a marine environment, yet restricted in terms of salinity such that brackish conditions prevailed during deposition of the sortehat formation. conclusions the age of the sortehat formation is aalenian to ?early bajocian. this age assignment is based on comparison of the three palynomorph assemblage zones (the botryococcus assemblage zone, the nannoceratopsis gracilis – nannoceratopsis senex assemblage zone and the sentusidinium pelionense assemblage zone) with palynomorph assemblages in the north atlantic region that have been recorded either from sections dated directly by ammonites or from sections that can be reliably correlated to successions with good ammonite control. no major changes have been detected in the miospore and microplankton assemblages across the lower boundary of the sortehat formation. this confirms the suggestion by underhill & partington (1994) that the boundary between the ostreaelv formation and the sortehat formation does not represent a biostratigraphically significant hiatal surface. the change in depositional environment from the sandy, shallow marine deposits of the ostreaelv formation to the mudstones of the sortehat formation records a marine drowning event. this sharp lithological boundary is broadly coincident with an influx of botryococcus (the botryococcus assemblage zone). the abundant occurrence of the freshwater green alga botryococcus together with marine dinoflagellate cysts is suggested to represent an allochthonous accumulation of botryococcus in a marine environment caused by seawards transport of material from an inland freshwater environment during the marine transgression. in general, there is an increase in the marine signal passing up-section from the botryococcus assemblage zone, as testified by the diversification of marine microplankton in the nannoceratopsis gracilis – nannoceratopsis senex assemblage zone associated with a decrease in the proportion of spores and the disappearance of botryococcus. in the uppermost part of the sortehat formation (the upper part of the sentusidinium pelionense assemblage zone), the palynomorph assemblages become poor in diversity and richness, suggesting more stressed conditions, but dinoflagellate cysts continue into the sandy pelion formation. acknowledgements this study was supported by a grant from the danish energy research programme (efp-93, projects 1313/930010, -0017) to the geological survey of greenland. we are grateful to david j. batten, karen dybkjær, jon r. ineson, henrik nøhr-hansen, stefan piasecki and finn surlyk for discussion and critical reading of the manuscript and to henrik nøhr-hansen in particular for patient 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(eds): geology of the brent group. geological society special publication (london) 61, 203–212. woollam, r. & riding, j.b. 1983: dinoflagellate cyst zonation of the english jurassic. institute of geological sciences report 83/2, 42 pp. london: her majesty’s stationery office. 795 manuscript received 16 april 1997; revision accepted 18 january 1999. appendix 1: locality information and geological survey of greenland (ggu) numbers of samples used in this study albuen location. 70°34′10′′n, 22°38′54′′w. section datum. the boundary between the ostreaelv and the sortehat formations is the datum for the measured section and is located 440 m above sea level. sample numbers. ggu 397452–397498, 398501, 398507, 398509, 398513. sortehat borehole location. 70°53′50′′n, 22°49′56′′w. section datum. for the drill-core, an arbitrary datum was defined 80 m below the top of the borehole. the boundary between the ostreaelv and the sortehat formations occurs 27.6 m above this datum. sample numbers. ggu 303143-11–303143-80. lepidopteriselv location. 71°15′12′′n, 22°37′04′′w. section datum. the boundary between the ostreaelv and the sortehat formations, which forms the datum for the section, lies at c. 810 m above sea level. sample numbers. ggu 398151–398199. enhjørningen dal two sections were measured in enhjørningen dal; figure 7 is a composite of these sections. the boundary between the ostreaelv and the sortehat formations, at 418 m above sea level, is the datum for the composite section. west enhjørningen dal location. 71°26′53′′n, 23°25′41′′w; the northernmost of two small ravines. sample numbers. ggu 398203–398299, 398308–398372. east enhjørningen dal location. 71°26′14′′n, 23°21′52′′w; the northernmost of two small ravines. sample numbers. ggu 395616–395668, 398417–398448. pelion location. 71°28′n, 23°19′w. sample numbers. ggu 339702–339711. these samples were collected by lars stemmerik in 1990. 796 appendix 2: list of all recorded palynomorph taxa miospores: alisporites robustus nilsson 1958 apiculatisporites spp. araucariacites australis cookson 1947 baculatisporites comaumensis (cookson) potonié 1956 (plate 1, fig. 3) b. spp. bisaccates indeterminate callialasporites dampieri (balme) dev 1961 (plate 2, fig. 1) c. microvelatus schulz 1966 (plate 2, fig. 3) c. minus (tralau) guy 1971 c. segmentatus (balme) dev 1961 c. trilobatus (balme) dev 1961 c. turbatus (balme) schulz 1967 (plate 2, fig. 2) c. spp. calamospora tener (leschik) mädler 1964 camarozonasporites spp. cerebropollenites macroverrucosus (thiergart) schulz 1967 (plate 3, fig. 3) c. thiergartii schulz 1967 chasmatosporites apertus nilsson 1958 c. hians nilsson 1958 (plate 2, fig. 4) c. major nilsson 1958 (plate 2, fig. 6) chomotriletes minor (kedves) pocock 1970 (plate 4, fig. 1) cibotiumspora jurienensis (balme) filatoff 1975 conbaculatisporites mesozoicus klaus 1960 c. spp. concavissimisporites spp corollina torosus (reissinger) cornet & traverse 1975 (plate 3, fig. 2) c. spp. deltoidospora minor (couper) pocock 1970 d. spp. d. toralis (leschik) lund 1977 densoisporites scanicus tralau 1968 d. velatus weyland & krieger 1953 densosporites spp. eucommiidites troedsonii erdtman 1948 exesipollenites tumulus balme 1957 (plate 3, fig. 1) foraminisporis jurassicus schulz 1967 fungal spores iraqispora labrata singh 1964 ischyosporites crateris balme 1957 i. spp. (plate 1, fig. 6) i. variegatus (couper) schulz 1967 kekryphalospora distincta fenton & riding 1987 kraeuselisporites reissingeri (harris) morbey 1975 leptolepidites bossus (couper) schulz 1967 l. major couper 1958 l. spp. limbosporites lundbladii nilsson 1958 lycopodiacidites rugulatus (couper) schulz 1967 manumia delcourtii (pocock) dybkjær 1991 (plate 1, fig. 8) megaspore fragments monosaccate pollen monosulcites spp. murospora spp. neoraistrickia gristhorpensis (couper) tralau 1967 n. taylorii playford & dettmann 1965 n. spp. perinopollenites elatoides couper 1958 (plate 2, fig. 5) pinuspollenites minimus (couper) kemp 1970 (plate 3, fig. 6) polycingulatisporites circulus simoncsics & kedves 1961 polycingulatisporites triangularis (bolkhovitina) playford & dettmann 1965 quadraeculina anellaeformis malyavkina 1949 (plate 3, fig. 5) retitriletes austroclavatidites (cookson) döring et al. 1963 (plate 1, fig. 1) r. clavatoides (couper) döring et al. 1963 (plate 1, fig. 2) r. semimuris (danzé-corsin & laveine) mckellar 1974 r. spp. ricciisporites tuberculatus lundblad 1954 rogalskaisporites cicatricosus (rogalska) danzé-corsin & laveine 1963 sestrosporites pseudoalveolatus (couper) dettmann 1963 spheripollenites psilatus couper 1958 s. spp. (plate 3, fig. 4) s. subgranulatus couper 1958 staplinisporites caminus (balme) pocock 1970 (plate 1, fig. 5) s. spp. stereisporites stereoides (potonié & venitz) h.d. pflug in: thomson & pflug 1953 striatella jurassica mädler 1964b (plate 1, fig. 4) s. parva (li & shang) filatoff & price 1988 s. seebergensis mädler 1964b s. spp. striate spp. taeniasporites rhaeticus schulz 1967 t. spp. taurocusporites verrucatus schulz 1967 t. spp. tigrisporites spp. todisporites major couper 1958 t. minor couper 1958 triancoraesporites spp. trilete spp. tripartina variabilis malyavkina 1949 (plate 1, fig. 7) 797 798 uvaesporites puzzlei guy 1971 u. spp. vesicaspora fuscus (pautsch) morbey 1975 vitreisporites pallidus (reissinger) nilsson 1958 zebrasporites laevigatus (schulz) schulz 1967 phytoplankton: acritarch spp. andreedinium arcticum below 1987 beaumontella caminuspina (wall) below 1987 botryococcus spp. (plate 4, fig. 4) chytroeisphaeridia chytroeoides (sarjeant) downie & sarjeant 1965 cymatiosphaera spp. dinocyst spp. dissilodinium spp. (plate 6, figs 2, 4) dodekovia tegillata prauss 1989 foraminiferal linings (plate 4, fig. 3) fungal spp. kallosphaeridium spp. lecaniella foveata singh 1971 leiosphaeridia spp. limbicysta bjaerkei (smelror) macrae et al. 1996 mancodinium semitabulatum morgenroth 1970 (plate 6, fig. 1) m. spp. mendicodinium groenlandicum (pocock & sarjeant) davey 1979 (plate 6, fig. 3) m. reticulatum morgenroth 1970 m. spp. miscellaneous nannoceratopsis ambonis drugg 1978 (plate 5, fig. 3) n. gracilis alberti emend. van helden 1977 (plate 5, fig. 1) n. plegas drugg 1978 (plate 5, fig. 4) n. ridingii poulsen 1992 n. senex van helden 1977 (plate 5, fig. 2) n. spp. n. triangulata prauss 1987 n. triceras drugg 1978 pareodinia ceratophora deflandre 1947 p. halosa (filatoff) prauss 1989 (plate 4, fig. 6) p. spp. parvocysta barbata bjærke 1980 p. spp. phallocysta eumekes dörhöfer & davies 1980 p. thomasi smelror 1991 p. spp. pterospermella spp scrinocassis weberi gocht 1964 s. spp. sentusidinium pelionense fensome 1979 (plate 6, figs 5, 6) s. spp. susadinium scrofoides (dörhöfer & davies) below 1987 tasmanites spp. valensiella ovulum (deflandre) eisenack 1963 veryhachium collectum wall 1965 v. formosum stockmans & williere 1960 v. sortehatense fensome 1979 (plate 4, fig. 2) v. spp. wallodinium laganum feist-burkhardt & monteil 1994 (plate 4, fig. 5) 799 plates 1–6 800 plate 1 figs 1–4 and 6–8 are from the lepidopteriselv section, fig. 5 is from the borehole at sortehat, the type locality of the sortehat formation. scale bar is 10 microns. for each of the illustrated specimens, the efr (england finder reference) is given. fig. 1. retitriletes austroclavatidites. sample 398181, slide 3, efr e36. fig. 2. retitriletes clavatoides. sample 398181, slide 3, efr f50. fig. 3. baculatisporites comaumensis. sample 398158, slide 4, efr s26. fig. 4. striatella jurassica. sample 398158, slide 5, efr k271. fig. 5. staplinisporites caminus. sample 303143-26, slide 3, efr j51. fig. 6. ischyosporites sp. sample 398189, slide 4, efr m40. fig. 7. tripartina variabilis. sample 398189, slide 4, efr r43. fig. 8. manumia delcourtii. sample 398192, slide 4, efr w383. 801 1 2 4 3 5 6 7 8 802 plate 2 figs 1 and 4–6 are from the borehole at sortehat, figs 2 and 3 are from the lepidopteriselv section. scale bar is 10 microns. fig. 1. callialasporites dampieri. sample 303143-46, slide 3, efr f222. fig. 2. callialasporites turbatus. sample 398158, slide 4, efr p25. fig. 3. callialasporites microvelatus. sample 398158, slide 4, efr h382. fig. 4. chasmatosporites hians. sample 303143-46, slide 3, efr l42. fig. 5. perinopollenites elatoides. sample 303143-13, slide 3, efr n25. fig. 6. chasmatosporites major. sample 303143-46, slide 3, efr k43. 803 1 2 4 3 5 6 804 plate 3 figs 1, 2, 5 and 6 are from the lepidopteriselv section, figs 3 and 4 are from the borehole at sortehat. scale bar is 10 microns. fig. 1. exesipollenites tumulus. sample 398181, slide 3, efr j362. fig. 2. corollina torosus. sample 398158, slide 5, efr m33. fig. 3. cerebropollenites macroverrucosus. sample 303143-40, slide 3, efr f24. fig. 4. spheripollenites sp. sample 303143-46, slide 3, efr w20. fig. 5. quadraeculina anellaeformis. sample 398158, slide 4, efr s384. fig. 6. pinuspollenites minimus. sample 398158, slide 5, efr d522. 805 1 2 4 3 5 6 806 plate 4 figs 1, 2 and 6 are from the borehole at sortehat, figs 3–5 are from the lepidopteriselv section. scale bar is 10 microns. fig. 1. chomotriletes minor. sample 303143-46, slide 3, efr k46. fig. 2. veryhachium sortehatense. sample 303143-46, slide 3, efr h22. fig. 3. foraminiferal inner lining. sample 398158, slide 4, efr h31. fig. 4. botryococcus sp. sample 398189, slide 4, efr e492. fig. 5. wallodinium laganum. sample 398194, slide 4, efr z39. fig. 6. pareodinia halosa. sample 303143-26, slide 3, efr m51. 807 1 2 4 3 5 6 808 plate 5 fig. 1 is from the lepidopteriselv section, figs 2–4 are from the borehole at sortehat. scale bar is 10 microns. fig. 1. nannoceratopsis gracilis. sample 398158, slide 4, efr h54. fig. 2. nannoceratopsis senex. sample 303143-40, slide 3, efr s363. fig. 3. nannoceratopsis ambonis. sample 303143-46, slide 3, efr j40. fig. 4. nannoceratopsis plegas. sample 303143-40, slide 3, efr m482. 809 1 2 43 810 plate 6 figs 1, 2 and 4–6 are from the borehole at sortehat, fig. 3 is from the lepidopteriselv section. scale bar is 10 microns. fig. 1. mancodinium semitabulatum. sample 303143-46, slide 3, efr s264. fig. 2. dissilodinium sp. sample 303143-30, slide 3, efr k324. fig. 3. mendicodinium groenlandicum. sample 398158, slide 5, efr r39. fig. 4. dissilodinium sp. sample 303143-26, slide 3, efr c54. fig. 5. sentusidinium pelionense. sample 303143-20, slide 3, efr m51. fig. 6. sentusidinium pelionense. sample 303143-26, slide 3, efr e53. 811 1 2 3 4 5 6 albuen (a) lo w er ju ra ss ic m a to ar ci an o st re ae lv f or m at io n g ul e h or n fo rm at io n sk æ vd al m em be r n at ho rs t fj el d m em be r a lb ue n m b a s el is b je rg m em be r tr ef jo rd b je rg m b l. p lie ns ba ch ia n u pp er p lie ns ba ch ia n 439.00 438.50 434.00 432.00 431.50 428.00 425.90 416.50 411.50 402.00 392.00 379.40 377.50 375.50 369.50 367.50 366.20 365.00 363.00 361.30 359.00 297.80 295.40 294.20 292.20 290.75 290.20 287.80 287.00 284.70 283.40 273.80 271.00 267.40 264.27 264.25 256.50 253.40 250.50 249.00 247.40 246.00 241.00 240.00 238.20 238.00 236.80 231.70 229.80 229.50 229.40 229.00 222.50 215.50 213.00 211.00 341248 397452 405449 341247 405450 405451 405452 405453 405454 405455 405456 405457 341243 405458 405459 341245 405460 341241 405462 405464 405466 405434 405433 405432 405431 405430 405429 341236 405428 405427 405426 405425 341235 405424 405423 405422 341234 405421 405420 341233 405419 405418 405417 405416 405414 405413 405411 405410 405408 341232 405406 405405 405404 405403 405402 405401 1 ro ga lsk ai sp or ite s ci ca tr ic os us 2 co nb ac ul at isp or ite s m es oz oi cu s 3 ly co po di ac id ite s ru gu la tu s 4 d el to id os po ra s pp . 5 re tit ril et es c la va to id es 6 re tit ril et es s em im ur is 7 re tit ril et es s p. 8 ba cu la tis po rit es s p. 9 ti gr isp or ite s m ic ro ru gu la tu s 10 fo ra m in isp or is ju ra ss ic us 11 kr ae us el isp or ite s re iss in ge ri 12 tr ip ar tin a va ria bi lis 13 st er ei sp or ite s st er eo id es 14 d en so isp or ite s sc an ic us 15 re tit ril et es a us tr oc la va to id es 16 to di sp or ite s m in or 17 ke kr yp ha lo sp or a di st in ct a 18 ci bo tiu m sp or ite s ju rie ne ns is 19 st ria te lla p ar va 20 to di sp or ite s m aj or 21 st ria te lla ju ra ss ic a 22 le pt ol ep id ite s sp . 23 z eb ra sp or ite s in te rs cr ip tu s 24 st er ei sp or ite s an tiq ua sp or ite s 25 u va es po rit es a rg en te ae fo rm is 26 tr ile te s sp . 27 st ria te lla s pp . 28 st ria te lla s ee be rg en sis 29 ch om ot ril et es s p. 30 m eg as po re fr ag m en ts 31 an ap ic ul at isp or ite s sp . 32 an ap ic ul at isp or ite s te le ph or us 33 n eo ra ist ric ki a sp . 34 ta ur oc us po rit es v er ru ca tu s 35 d en so isp or ite s ve la tu s 36 sc ul pt isp or ite s au lo se ne ns is 37 u va es po rit es s p. 38 to di sp or ite s sp . 39 m an um ia d el co ur tii 40 is ch yo sp or ite s va rie ga tu s 41 is ch yo sp or ite s cr at er is 42 st ap lin isp or ite s ca m in us 43 po lyc in gu la tis po rit es tr ia ng ul ar is 44 se st ro sp or ite s ps eu do al ve ol at us 45 po lyc in gu la tis po rit es c irc ul us 46 n eo ra ist ric ki a ta ylo rii 47 le pt ol ep id ite s m aj or 48 li m bo sp or ite s lu nd bl ad ii 49 m ur os po ra s p. 50 ch as m at os po rit es h ia ns 51 ce re br op ol le ni te s th ie rg ar tii 52 ve sic as po ra fu sc us 53 q ua dr ae cu lin a an el la ef or m is 54 pe rin op ol le ni te s el at oi de s 55 pi nu sp ol le ni te s m in im us 56 bi sa cc at e sp p. 57 ce re br op ol le ni te s m ac ro ve rr uc os us 58 ch as m at os po rit es m aj or 59 ch as m at os po rit es a pe rt us 60 co ro llin a to ro su s 61 m on os ul ci te s pu nc ta tu s 62 ch as m at os po rit es s p. 63 ch as m at os po rit es e le ga ns 64 m on os ac ca te s pp . 65 vi tr ei sp or ite s pa llid us 66 ca llia la sp or ite s da m pi er i 67 ca llia la sp or ite s sp . 68 ar au ca ria ci te s au st ra lis 69 eu co m m iid ite s tr oe ds on ii 70 ce re br op ol le ni te s sp . 71 sp he rip ol le ni te s su bg ra nu la tu s 72 ca m pe ni a sp . 73 ca llia la sp or ite s m ic ro ve la tu s 74 ca llia la sp or ite s m in us 75 ex es ip ol le ni te s tu m ul us 76 ca llia la sp or ite s tr ilo ba tu s 77 ca llia la sp or ite s tu rb at us 78 eu co m m iid ite s m aj or 79 ri cc iis po rit es tu be rc ul at us 80 ta en ia sp or ite s sp . 81 st ria te s pp . ? ? ? r r ? ? r r r r r r r r alphabetical species list 31 anapiculatisporites sp. 32 anapiculatisporites telephorus 68 araucariacites australis 8 baculatisporites sp. 56 bisaccate spp. 66 callialasporites dampieri 73 callialasporites microvelatus 74 callialasporites minus 67 callialasporites sp. 76 callialasporites trilobatus 77 callialasporites turbatus 72 campenia sp. 57 cerebropollenites macroverrucosus 70 cerebropollenites sp. 51 cerebropollenites thiergartii 59 chasmatosporites apertus 63 chasmatosporites elegans 50 chasmatosporites hians 58 chasmatosporites major 62 chasmatosporites sp. 29 chomotriletes sp. 18 cibotiumsporites jurienensis 2 conbaculatisporites mesozoicus 60 corollina torosus 4 deltoidospora spp. 14 densoisporites scanicus 35 densoisporites velatus 78 eucommiidites major 69 eucommiidites troedsonii 75 exesipollenites tumulus 10 foraminisporis jurassicus 41 ischyosporites crateris 40 ischyosporites variegatus 17 kekryphalospora distincta 11 kraeuselisporites reissingeri 47 leptolepidites major 22 leptolepidites sp. 48 limbosporites lundbladii 3 lycopodiacidites rugulatus 39 manumiadel courtii 30 megaspore fragments 64 monosaccate spp. 61 monosulcites punctatus 49 murospora sp. 33 neoraistrickia sp. 46 neoraistrickia taylorii 54 perinopollenites elatoides 55 pinuspollenites minimus 45 polycingulatisporites circulus 43 polycingulatisporites triangularis 53 quadraeculinaanellae formis 15 retitriletes austroclavatoides 5 retitriletes clavatoides 6 retitriletes semimuris 7 retitriletes sp. 79 ricciisporites tuberculatus 1 rogalskaisporites cicatricosus 36 sculptisporites aulosenensis 44 sestrosporites pseudoalveolatus 71 spheripollenites subgranulatus 42 staplinisporites caminus 24 stereisporites antiquasporites 13 stereisporites stereoides 81 striate sp. 21 striatella jurassica 19 striatella parva 28 striatella seebergensis 27 striatella sp. 80 taeniasporites sp. 34 taurocusporites verrucatus 9 tigrisporites microrugulatus 20 todisporites major 16 todisporites minor 38 todisporites sp. 26 triletes sp. 12 tripartina variabilis 25 uvaesporites argenteaeformis 37 uvaesporites sp. 52 vesicaspora fuscus 65 vitreisporites pallidus 23 zebrasporites interscriptus 400 380 360 300 280 260 240 220 420 interval not sampled sy st em st ag e pa ly no lo gi ca l a ss em bl . z on es li th os tr at ig ra ph y (m ) sa m pl e he ig ht sa m pl e nu m be r 7 6 5 4 3 2 1 uncertain determination very rare rare few common abundant ? r fig. 4a. terrestrial palynomorph distribution chart for the gule horn and ostreaelv formations at albuen (for location, see fig. 1). m, middle jurassic; a, aalenian; as, astartekløft member. albuen (b) 439.00 438.50 434.00 432.00 431.50 428.00 425.90 416.50 411.50 402.00 392.00 379.40 377.50 375.50 369.50 367.50 366.20 365.00 363.00 361.30 359.00 292.20 287.80 287.00 284.70 283.40 273.80 271.00 267.40 264.27 264.25 256.50 253.40 250.50 249.00 247.40 246.00 241.00 240.00 238.20 238.00 236.80 231.70 229.80 229.50 229.40 229.00 222.50 215.50 213.00 211.00 341248 397452 405449 341247 405450 405451 405452 405453 405454 405455 405456 405457 341243 405458 405459 341245 405460 341241 405462 405464 405466 405431 341236 405428 405427 405426 405425 341235 405424 405423 405422 341234 405421 405420 341233 405419 405418 405417 405416 405414 405413 405411 405410 405408 341232 405406 405405 405404 405403 405402 405401 1 n an no ce ra to ps is se ne x 2 n an no ce ra to ps is sp . 3 n an no ce ra to ps is gr ac ilis 4 m en di co di ni um r et ic ul at um 5 be au m on te lla c am in us pi na 6 n an no ce ra to ps is tr ia ng ul at a 7 n an no ce ra to ps is pl eg as 8 m an co di ni um s em ita bu la tu m 9 pa rv oc ys ta b ar ba ta 10 m en di co di ni um g ro en la nd ic um 11 be au m on te lla d el ic at a 12 m en di co di ni um s p. 13 lu eh nd ea s pi no sa 14 d in oc ys t sp . 15 va lva eo di ni um a rm at um 16 va lva eo di ni um s pp . 17 n an no ce ra to ps is am bo ni s 18 d iss ilio di ni um s p. 19 pa rv oc ys ta s p. 20 ph al lo cy st a eu m ek es 21 n an no ce ra to ps is tr ic er as 22 pa re od in ia h al os a 23 ka llo sp ha er id iu m s p. 24 sc rin io ca ss is sp . 25 ph al lo cy st a el on ga ta 26 w al lo di ni um s pp . 27 a cr ita rc h sp p. 28 le io fu sa ju ra ss ic a 29 le io sp ha er id ia s pp . 30 m ic rh ys tr id iu m ly m en sis 31 m ic rh ys tr id iu m s pp . 32 ve ry ha ch iu m fo rm os um 33 li m bi cy st a bj ae rk ei 34 ve ry ha ch iu m c ol le ct um 35 te tr ap or in a co m pr es sa 36 le ca ni el la fo ve at a 37 ve ry ha ch iu m r ed uc tu m 38 ve ry ha ch iu m ir re gu la re 39 m ic rh ys tr id iu m in tr om itt um 40 m ic rh ys tr id iu m w at to ne ns e 41 m ic rh ys tr id iu m fr ag ile 42 ve ry ha ch iu m tr isp in os um 43 m ic rh ys tr id iu m s te lla tu m 44 cy m at io sp ha er a sp . 45 be au m on te lla s p. 46 pt er os pe rm el la s pp . 47 bo tr yo co cc us s pp . 48 ta sm an ite s sp . 49 le ca ni el la s pp . 50 fu ng al s po re s 51 ce lyp hu s sp p. 52 h ap lo ph ra gm oi de s sp p. 53 m is ce lla ne ou s sp p. ? ? ? ? ? ? ? alphabetical species list 27 acritarch spp. 5 beaumontella caminuspina 11 beaumontella delicata 45 beaumontella sp. 47 botryococcus spp. 51 celyphus spp. 44 cymatiosphaera sp. 14 dinocyst sp. 18 dissiliodinium sp. 50 fungal spores 52 haplophragmoides spp. 23 kallosphaeridium sp. 36 lecaniella foveata 49 lecaniella spp. 28 leiofusa jurassica 29 leiosphaeridia spp. 33 limbicysta bjaerkei 13 luehndea spinosa 53 miscellaneous spp. 8 mancodinium semitabulatum 10 mendicodinium groenlandicum 4 mendicodinium reticulatum 12 mendicodinium sp. 41 micrhystridium fragile 39 micrhystridium intromittum 30 micrhystridium lymensis 31 micrhystridium spp. 43 micrhystridium stellatum 40 micrhystridium wattonense 17 nannoceratopsis ambonis 3 nannoceratopsis gracilis 7 nannoceratopsis plegas 1 nannoceratopsis senex 2 nannoceratopsis sp. 6 nannoceratopsis triangulata 21 nannoceratopsis triceras 22 pareodinia halosa 9 parvocysta barbata 19 parvocysta sp. 25 phallocysta elongata 20 phallocysta eumekes 46 pterospermella spp. 24 scriniocassis sp. 48 tasmanites sp. 35 tetraporina compressa 15 valvaeodinium armatum 16 valvaeodinium spp. 34 veryhachium collectum 32 veryhachium formosum 38 veryhachium irregulare 37 veryhachium reductum 42 veryhachium trispinosum 26 wallodinium spp. sy st em st ag e pa ly no lo gi ca l a ss em bl . z on es li th os tr at ig ra ph y (m ) sa m pl e he ig ht sa m pl e nu m be r interval not sampled 1 2 3 4 5 6 7 lo w er ju ra ss ic m a to ar ci an o st re ae lv f or m at io n g ul e h or n fo rm at io n sk æ vd al m em be r n at ho rs t fj el d m em be r a lb ue n m b a s el is b je rg m em be r tr ef jo rd b je rg m b l. p lie ns ba ch ia n u pp er p lie ns ba ch ia n 400 380 360 300 280 260 240 220 420 uncertain determination very rare rare few common abundant ? r fig. 4b. marine palynomorph distribution chart for the gule horn and ostreaelv formations at albuen (for location, see fig. 1). m, middle jurassic; a, aalenian; as, astartekløft member. lepidopteriselv lo w er ju ra ss ic u pp er p lie ns ba ch ia n g ul e h or n fo rm at io n el is b je rg m em be r 700.00 695.00 685.00 679.00 674.00 654.00 653.00 648.00 642.00 638.00 139146 139145 139144 139143 139142 139141 139140 139139 139138 139137 1 ba cu la tis po rit es s p. 2 d el to id os po ra s pp . 3 ca la m os po ra te ne r 4 to di sp or ite s m aj or 5 ro ga lsk ai sp or ite s ci ca tr ic os us 6 st ria te lla p ar va 7 re tit ril et es a us tr oc la va to id es 8 re tit ril et es c la va to id es 9 n eo ra ist ric ki a sp . 10 st er ei sp or ite s st er eo id es 11 re tit ril et es s em im ur is 12 kr ae us el isp or ite s re iss in ge rii 13 re tit ril et es s p. 14 ti gr isp or ite s m ic ro ru gu la tu s 15 ap ic ul at isp or ite s pa rv isp in os us 16 co nb ac ul at isp or ite s m es oz oi cu s 17 d en so isp or ite s sc an ic us 18 ly co po di ac id ite s ru gu la tu s 19 ira qu isp or a sp . 20 ti gr isp or ite s sp . 21 tr ip ar tin a va ria bi lis 22 ke kr yp ha lo sp or a di st in ct a 23 an ap ic ul at isp or ite s sp . 24 m ar at tii sp or ite s sc ab ra tu s 25 d el to id os po ra m in or 26 an nu lis po ra fo llic ul os a 27 ap ic ul at isp or ite s sp . 28 st ria te lla ju ra ss ic a 29 d en so sp or ite s sp . 30 d en so sp or ite s va ria bi lis 31 ca m ar oz on os po rit es r ud is 32 m ur os po ra s p. 33 q ua dr ae cu lin a an el la ef or m is 34 bi sa cc at e sp p. 35 pe rin op ol le ni te s el at oi de s 36 ch as m at os po rit es h ia ns 37 pi nu sp ol le ni te s m in im us 38 ch as m at os po rit es a pe rt us 39 ce re br op ol le ni te s th ie rg ar tii 40 co ro llin a to ro su s 41 m on os ul ci te s pu nc ta tu s 42 ce re br op ol le ni te s m ac ro ve rr uc os us 43 ve sic as po ra fu sc us 44 ch as m at os po rit es m aj or 45 ar au ca ria ci te s au st ra lis 46 sp he rip ol le ni te s ps ila tu s 47 ca llia la sp or ite s tu rb at us 48 ca llia la sp or ite s m in us 49 eu co m m iid ite s tr oe ds on ii 50 m on os ac ca te s pp . 51 co ro llin a sp p. 52 vi tta tin a sp . 53 m en di co di ni um r et ic ul at um 54 m an co di ni um s em ita bu la tu m 55 n an no ce ra to ps is se ne x 56 n an no ce ra to ps is tr ia ng ul at a 57 n an no ce ra to ps is gr ac ilis 58 n an no ce ra to ps is sp . 59 m ic rh ys tr id iu m in tr om itt um 60 ve ry ha ch iu m tr isu lc um 61 m ic rh ys tr id iu m fr ag ile 62 le ca ni el la s pp . 63 m ic rh ys tr id iu m ly m en sis 64 ve ry ha ch iu m s p. 65 bo tr yo co cc us s pp . 66 ta sm an ite s sp . ? r r ? r ? r ? alphabetical species list 23 anapiculatisporites sp. 26 annulispora folliculosa 15 apiculatisporites parvispinosus 27 apiculatisporites sp. 45 araucariacites australis 1 baculatisporites sp. 34 bisaccate spp. 65 botryococcus spp. 3 calamospora tener 48 callialasporites minus 47 callialasporites turbatus 31 camarozonosporites rudis 42 cerebropollenites macroverrucosus 39 cerebropollenites thiergartii 38 chasmatosporites apertus 36 chasmatosporites hians 44 chasmatosporites major 16 conbaculatisporites mesozoicus 51 corollina sp. 40 corollina torosus 25 deltoidospora minor 2 deltoidospora spp. 17 densoisporites scanicus 29 densosporites sp. 30 densosporites variabilis 49 eucommiidites troedsonii 19 iraquispora sp. 22 kekryphalospora distincta 12 kraeuselisporites reissingerii 62 lecaniella spp. 18 lycopodiacidites rugulatus 54 mancodinium semitabulatum 24 marattiisporites scabratus 53 mendicodinium reticulatum 61 micrhystridium fragile 59 micrhystridium intromittum 63 micrhystridium lymensis 50 monosaccate spp. 41 monosulcites punctatus 32 murospora sp. 57 nannoceratopsis gracilis 55 nannoceratopsis senex 58 nannoceratopsis sp. 56 nannoceratopsis triangulata 9 neoraistrickia sp. 35 perinopollenites elatoides 37 pinuspollenites minimus 33 quadraeculina anellaeformis 7 retitriletes austroclavatoides 8 retitriletes clavatoides 11 retitriletes semimuris 13 retitriletes sp. 5 rogalskaisporites cicatricosus 46 spheripollenites psilatus 10 stereisporites stereoides 28 striatella jurassica 6 striatella parva 66 tasmanites sp. 14 tigrisporites microrugulatus 20 tigrisporites sp. 4 todisporites major 21 tripartina variabilis 64 veryhachium sp. 60 veryhachium trisulcum 43 vesicaspora fuscus 52 vittatina sp. sy st em st ag e pa ly no lo gi ca l a ss em bl . z on es li th os tr at ig ra ph y (m ) sa m pl e he ig ht sa m pl e nu m be r 700 675 650 3 2 uncertain determination very rare rare few common abundant ? r fig. 13. palynomorph distribution chart for the gule horn formation (elis bjerg member) at lepidopteriselv (for location, see fig. 1). geological survey of denmark and greenland bulletin 4, , pp. 33-36 33 the first dna-based methods for direct quantification of soil protozoa, and a dna-based quantification method to describe the spread of phenanthrene-degrading bacteria in soil and freshwater aquifers, have recently been developed at the biopro research centre at the geological survey of denmark and greenland (geus). well-known genes for phenoxyalcanoic acid degradation have been used to monitor the in situ degradation of phenoxyalcanoic acid pesticides. studies have been initiated on the short-lived mrna molecules that are expected to provide a shortcut to the understanding of low, yet important, microbial activity in geological samples. this article reviews recent developments in techniques based on analysis of nucleic acids from soils and aquifers. analytical work has been carried out mainly on soil samples from a former asphalt production plant at ringe (fig. 1). the ringe plant constitutes one of the most polluted industrial sites in denmark, and is a priority site of studies by the biopro research centre. although rich in carbon, the ringe subsoil is an oligotrophic environment due to the high content of polycyclic aromatic hydrocarbons (pah). this is an environment where the supply of nutrients to microorganisms is low, leading to slow growth, low total numbers of microorganisms and small cells. to study microbial communities of oligotrophic environments, analytical methods with low detection limits are needed. until recently, microorganisms of natural environments were mainly studied by cultivation-dependent methods. however, microorganisms that can be cultured on agar plates are now known to represent only a small fraction of the total microbial community. modern methods, therefore, need to be based on the detection of biomolecules in the microorganisms rather than being dependent on growth of the microorganisms. the best available techniques are based on dna and rna molecules (fig. 2), which due to their high level of resolution allow closely related organisms or functional genes to be distinguished. in the following review, examples are given of applications of these nucleic acid based methods. direct analysis of microbial populations in soil and freshwater aquifers using nucleic acid based techniques carsten s. jacobsen, julia r. de lipthay, mikkel bender, line fredslund, anders r. johnsen and kaare johnsen fig. 1. location of the former asphalt plant at ringe, one of the most polluted industrial sites in denmark. the biopro research centre at geus monitored in situ degradation of phenoxyalcanoic acid pesticides. the pits, sampling sites and wells were used to determine the distribution of contamination of tar (pale red) and kerosene (green) in the subsoil. a field injection experiment, using a mixture of six different herbicides was carried out at a shallow sandy aquifer at vejen. modified from rosenbom et al. (2000). geological survey of denmark and greenland bulletin 4, 33–36 (2004) © geus, 2004 dna analysis is not limited to living organisms unlike other microbial analysis tools, dna analyses may be carried out on both living and dead material, as long as the dna has not been degraded. for analysis of dna in soils, a major problem is the presence of humic substances, that even at low concentrations interfere with the enzymes used in the dna amplification process. all research groups working with soil nucleic acids experience problems with amplification of dna when soils contain much humic material. one way to solve this problem is to use selective purification methods in which single-stranded dna is selectively purified from the soil using a molecular ‘fishing rod’ equipped with the complementary dna strand. after extracting the gene of interest, it is possible to multiply its numbers using the ‘polymerase chain reaction’ (pcr; saiki et al. 1985). pcr is an exponential reaction in which a single dna strand can, in principle, produce four million identical copies by 25 cycles of multiplication. in pcr, two small dna sequences, corresponding to two regions on the gene, are selected as priming sites for two segments of complementary dna. these small dna pieces are designated ‘primers’. the primers serve as the target of the pcr, and lead to the formation of a large number of dna molecules identical to the original gene. the primers are constructed by consulting dna sequence databases on the internet. since these databases are very comprehensive, it is possible to construct primer sets that are specific for the desired taxonomic or functional groups. this technique has, for example, enabled forensic experts to produce enough dna molecules to determine whether the genomic fingerprint of a person matches that of a bloodstain at the scene of a crime. we have used this technique widely for the analysis of microorganisms in the environment, and the detection limit is less than 40 cells in a sample (jacobsen 1995). the challenge is now not only to detect, but also to quantify the dna from very few cells in soil and freshwater aquifers. quantitative dna techniques for the enumeration of soil flagellates and bacteria as a consequence of the exponential nature of pcr, it is an excellent technique for the detection of specific dna sequences. on the other hand, it is not quantitative, and small differences in the efficiency of the reaction affect the final number of dna copies. several methods have been proposed to resolve this problem. one possibility is using the principle of ‘most probable number’ (mpn) estimates, where the dna template is serially diluted in several replicate reactions. by looking at which dilutions of template dna that give a product from the pcr, and in how many of the replicate reactions, an estimate of the original number of genes in the sample can be made. this approach has been used to develop the first successful molecular detection and quantification of protozoa in soil (fredslund et al. 2001). this new technique represents a breakthrough in reliable enumeration of soil protozoa, since these often small and amoeboid organisms are difficult to enumerate using microscopy techniques. traditionally, the enumeration was based on growth-dependent techniques, where the cultivable fraction of the total protozoan populations was not known. in our study, a part of the 18s rdna of the common soil flagellate heteromita globosa was sequenced and pcr primers for this gene were developed. in a sterilised soil at the ringe asphalt production plant, the population dynamics of this flagellate and the phenanthrene-degrading bacterium pseudomonas putida ous82 were quantified using both growth-dependent techniques and the mpn-pcr assay (fredslund et al. 2001). alternative methods for quantification of dna are the real-time pcr and the competitive pcr methods. in realtime pcr, the dna multiplication is monitored on-line using a combination of fluorescent dna stains and fibre optics coupled to a computer. geus has recently received funding from the danish natural science research council to implement this technique. competitive pcr makes use of an internal standard in the form of a similar, but shorter dna molecule, which is recognised and hence amplified by the same primer set as the template dna (the dna that needs to be quantified). the 34 fig. 2. all information regarding cell function and structure is contained as a genetic code in the cell dna. each gene encodes a specific function by dictating the synthesis of a specific protein. before proteins are synthesised, the genes are copied (transcribed) into messenger rna (mrna). after transcription, the mrna is translated into protein by protein-synthesising machinery called ribosomes. the sequence of bases in the mrna, copied after the base sequence in the gene, determines the structure and function of the protein. internal standard is added to the reaction mixtures in decreasing amounts and competes with the template dna for amplification. thus, the amount of product from the internal standard is inversely related to the initial amount of the template dna. competitive pcr exploits the highly sensitive nature of the pcr process, while using an internal standard to bypass the quantification problems inherent in the amplification reaction (johnsen et al. 1999). microbial changes in aquifers contaminated with phenoxyalcanoic acid herbicides this section focuses on specific genes of interest rather than on organisms. phenoxyalcanoic acid herbicides are extensively used in agriculture, and include compounds such as mecoprop (mcpp), 2,4-dichlorophenoxyacetic acid (2,4-d) and dichlorprop. a common pathway for 2,4-d degradation has been determined for the bacterial strain ralstonia eutropha jmp134, and the catabolic genes (tfd) encoding the specific enzymes have been identified (fig. 3; don et al. 1985). the in situ adaptation processes of the indigenous microorganisms when exposed to these herbicides have been investigated by studying the impact on a microbial community in a freshwater aquifer. a field injection experiment was carried out at a shallow sandy aquifer at vejen, denmark (fig. 1). during a seven-month period, a mixture of six different herbicides, including mcpp and dichlorprop, was continuously injected into the aquifer creating a contaminant plume (broholm et al. 2000). sediment and groundwater samples from herbicideexposed (1 and 2) and non-exposed (nx) sites (fig. 4a) were collected, and the impact on microbial community structure and function was studied (de lipthay et al. 2000). laboratory incubations demonstrated that sediment samples collected inside the contaminant plume had acquired a significantly increased capacity for herbicide mineralisation compared to samples from non-exposed sites (fig. 4a). thus, the in situ exposure to herbicides resulted in microbial communities that were better adapted to the degradation of phenoxyalcanoic acids. this was further demonstrated by greatly increased populations of pesticide degraders inside the pesticide plume, both when enumerated by cultivation, and when quantified by the number of pesticide genes (tfdabc) detected by pcr methods (fig. 4b, treatment 1 and 2). pesticide degraders and their tfd genes were undetectable outside the plume (fig. 4b, treatment nx). the most likely 35 fig. 3. pathway for degradation of 2,4-d as elucidated in the bacterial strain ralstonia eutropha jmp134. the tfda gene encodes a 2,4-d dioxygenase, tfdb encodes a 2,4-dichlorophenol hydroxylase, and tfdc encodes a chlorocatechol 1,2-dioxygenase. the 2,4-dichloromuconic acid generated by the activity of the tfdc gene product is further transformed to intermediates of the tricarboxylic acid cycle by the activity of other tfd gene products. fig. 4. a: mineralisation of the phenoxyalcanoic acid herbicides 2,4-d (red) and mcpp (blue) in laboratory incubations of sediment samples from herbicide exposed (1, 2) and non-exposed (nx) sites of the vejen aquifer. b: effect of in situ herbicide exposure in two exposed (1, 2) sites on microbial biomass of 2,4-d and mcpp degraders, and on the presence of tfda, tfdb and tfdc genes. data show that indigenous microbial communities carry the tfd genes and are capable of degrading phenoxyalcanoic acids. explanation is that microorganisms carrying the tfd genes had a selective advantage in that they could make use of the pesticides as sources of carbon and energy. the data suggest that natural attenuation is a likely procedure for clean up of this group of herbicide compounds when originating from pointsource contaminations. analysis of microbial activity applying mrna techniques the presence or absence of specific microorganisms may be determined by use of cultivation or dna-based techniques, although these methods give no information as to whether the organisms are actually active in the environment. microbial activity may be measured in several ways. the first sign of activity in microbial cells is the synthesis of messenger-rna (mrna; fig. 2). these molecules have half-lives of only a few minutes, and the detection of mrna thus ensures that the genes of interest are actually expressed at the time of sampling. another approach is to detect the activity of the gene products – the enzymes. however, the longevity of enzyme activities is variable. a third approach is to measure the target molecules of the enzymes, i.e. the pollutant molecules. by use of analytical chemical methods such as gas and liquid chromatography, the concentration of target molecules may be measured, and the dissipation of pollutants indicates microbial activity. a major topic of future studies in microbial ecology will be the assessment of microbial activity by the application of mrna techniques to answer which microorganisms are active, and under which conditions their genes are expressed. the two most commonly used techniques for detection of mrna (fig. 2) are reverse transcription polymerase chain reaction (rt-pcr) and rna-rna hybridisation. in rt-pcr, the first step is a reverse transcription process the conversion of mrna into copy dna (cdna). reverse transcription requires a small dna primer to bind to the mrna in order to initiate synthesis of cdna. thus, specific mrnas can be amplified by using sequence-specific primers in the rt-pcr reaction. following synthesis of cdna, a normal pcr is carried out to multiply the cdna, and the resulting pcr products are detected by usual dna detection techniques. direct rna-rna hybridisation analyses exploit a completely different principle; this is directly quantitative but lacks the sensitivity of rt-pcr. first, the total content of mrna is extracted. then the mrna of interest is identified by binding of a specific rna probe with a sequence complementary to the mrna gene of interest (hybridisation). by using a ‘radiolabelled’ probe, the final quantification of the mrna of interest is easily done by determining the amount of ‘radiolabelled’ bound to the mrna. a study using freshwater samples artificially contaminated with the herbicide 2,4-d, revealed a significant increase in the amount of tfda mrna, as measured by hybridisation of rna extracts using a tfda specific probe. this demonstrates that the 2,4-d degraders in the freshwater samples were actively degrading the 2,4-d (fig. 5). transcription of tfda was, however, transient and the degradation of 2,4-d continued although mrna levels dropped. these observations illustrate that the herbicide-degrading enzymes encoded by the mrna last longer in the cells than the mrna ‘signal’ itself. references broholm, m.m., rügge, k., tuxen, n., mosbæk, h., & bjerg, p.l. 2000: migration and degradation of pesticides in an aerobic groundwater aquifer: field injection experiments. in: bjerg, p.l., engesgaard, p. & krom, t.d. (eds): proceedings of the international conference on groundwater research, copenhagen, 169–170. rotterdam: balkema. de lipthay, j.r., johnsen, k., aamand, j., tuxen, n., albrechtsen, h.-j. & bjerg, p.l. 2000: continuous exposure of pesticides in an aquifer changes microbial biomass, diversity and degradation potential. in: bjerg, p.l., engesgaard, p. & krom, t.d. (eds): proceedings of the international conference on groundwater research, copenhagen, 157–158. rotterdam: balkema. don, r.h., weightman, a.j., knackmuss, h.-j. & timmis, k.n. 1985: transposon mutagenesis and cloning analysis of the pathway for degradation of 2,4-dichlorophenoxyacetic acid and 3-chlorobenzoate in alcaligenes eutrophus jmp134(pjp4). journal of bacteriology 161(1), 85–90. fredslund, l., ekelund, f., jacobsen, c.s. & johnsen, k. 2001: development and application of a most probable number-pcr assay to quantify flagellate populations in soil samples. applied and environmental microbiology 67(4), 1613–1618. jacobsen, c.s. 1995: rapid microscale detection of specific bacterial dna in soil using magnetic capture-hybridization and polymerase chain reaction amplification assay (mch-pcr). applied and environmental microbiology 61(9), 3347–3352. johnsen, k., enger, ø., jacobsen, c.s., thirup, l. & torsvik, v. 1999: quantitative selective pcr of 16s ribosomal dna correlates well with selective agar plating in describing population dynamics of indigenous pseudomonas spp. in soil hot spots. applied and environmental microbiology 65(5), 1786–1789. rosenbom, a.e., klint, k.e.s., fredericia, j., springer, n. & andersen, g. 2000: pore-to-core scale-up studies of the transport properties of organic pollutants with natural attenuation. danmarks og grønlands geologiske undersøgelse rapport 2000/79, 39 pp. saiki, r.k., scharf, s., faloona, f., mullis, k.b., horn, g.t., erlich, h.a. & arnheim, n. 1985: enzymatic amplification of ß-globin genomic sequences and restriction site analysis for diagnosis of sickle cell anemia. science 230, 1350–1354. 36 authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: csj@geus.dk << /ascii85encodepages false /allowtransparency false /autopositionepsfiles true /autorotatepages /all /binding /left /calgrayprofile (dot gain 20%) /calrgbprofile (srgb iec61966-2.1) 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<feff4f7f752890194e9b8a2d5b9a5efa7acb76840020005000440046002065874ef65305542b8f039ad876845f7150cf89e367905ea6ff0c4fbf65bc63d066075217537054c18cea3002005000440046002065874ef653ef4ee54f7f75280020004100630072006f0062006100740020548c002000520065006100640065007200200035002e0030002053ca66f465b07248672c4f86958b555f3002> >> >> setdistillerparams << /hwresolution [2400 2400] /pagesize [595.000 842.000] >> setpagedevice geological survey of denmark and greenland bulletin 26, 2012,13-16 13 geology of the femern bælt area between denmark and germany emma sheldon, peter gravesen and henrik nøhr-hansen geological and geotechnical investigations in the femern bælt area were undertaken from 1995 to 2010 (rambøll arup jv 2011) in preparation for the fixed link between lolland in denmark and fehmarn in germany. as a result, new data have been acquired on the stratigraphy and distribution of the deposits and the major structures and tectonic influence on the layers close to the surface. previous investigations of cretaceous–palaeogene deposits on southern lolland (fig.  1) were limited due to lack of outcrops and borehole data. two deep boreholes and geophysical surveys (1952– 1953) revealed: (1) the presence of a salt diapir at rødbyhavn, (2) upper maastrichtian chalk 29–143 m below quaternary deposits and (3) an erosional window in the palaeogene cover. boreholes to the east of rødbyhavn (1992–1994) revealed the sediment distribution on southern lolland and showed that cretaceous and palaeogene deposits are cut by several nw–se-orientated faults. this paper presents a summary of lithostratigraphic and biostratigraphic investigations and a brief description of the geological development in the area. methods a multidisciplinary biostratigraphic study was undertaken of 170 samples from 46 wells drilled in 2009–2010 in the femern bælt region. the wells were drilled to depths of 50– 100 m and were fully cored in the pre-quaternary deposits. one to ten samples from each well were selected for biostratigraphic analysis by rambøll arup jv. nannofossil analysis was carried out on all samples. supplementary analyses were carried out using microfossils and dinoflagellate cysts (dinocysts) when necessary. nannofossils are particularly useful for dating chalk and clay, microfossils for chalk, sand and clay and dinocysts for clay. the use of three fossil groups allowed for well-constrained and reliable dating and formation identification (sheldon & nøhr-hansen 2010). in addition, 13 samples from dgu core 241.213, a water well from 2011 on lolland, were analysed. biostratigraphy and lithostratigraphy a chronostratigraphic, biostratigraphic and lithostratigraphic correlation is presented in fig. 2. the lithologies and selected marker fossils reported in this article are as follows (see also fig. 3): tor formation equivalent (campanian–maastrichtian) – the tor formation of the north sea was originally described by deegan & scull (1977) as homogeneous white or grey chalk, bioturbated pelagic chalk, laminated chalk and redeposited chalk. the femern bælt area chalk is slightly indurated, with <5% flint nodules. biostratigraphic marker fossils: tranolithus orionatus, reinhardtites levis, nephrolithus frequens, arkhangelskiella maastrichtiana, eiffellithus turriseiffelii, prediscosphaera spinosa, cribrosphaerella daniae (nannofossils), bolivinoides laevigatus, praebulimina levis, reusella szajnochae szajnochae, bolivina incrassata, stensioeina pommerana (foraminifera), cannosphaeropsis utinensis, montanarocysta aemiliana and xenascus wetzelii (dinocysts). danian limestone (lower–middle danian) – comprises fossiliferous limestone, rich in echinoderm, bryozoan and bivalve fragments and is recognised as a thin layer on lolland (dgu core 241.213). biostratigraphic marker fossils: prinsius tenuiculus, prinsius dimorphosus, chiasmolithus danicus, cruciplacolithus tenuis, neochiastozygus modestus (nannofossils), eoglobigefig. 1. a: map of denmark and northern germany showing the proposed location of the fixed link across the femern bælt. b: geological map of the femern bælt region (after håkansson & pedersen 1992). lolland fehmarn femern bælt rødbyhavn eocene paleocene (without danian) danian upper cretaceous ringkøbing–fyn high 100 km 14°e 57°n 55° swedendenmark germany german basin danish basin a b dgu 241.213 norway © 2012 geus. geological survey of denmark and greenland bulletin 26, 13–16. open access: www.geus.dk/publications/bull 1414 rina edita, globoconusa daubjergensis and cibicidoides succedens (foraminifera). lellinge greensand (lower selandian) – in dgu core 241.213, lellinge greensand (originally described by gry 1935) comprises olive-grey, glauconitic, sandy, hard limestone intercalated with fine-grained clay with brown phosphate nodules and pyrite. lellinge greensand is also encountered in a number of femern bælt boreholes. biostratigraphic marker fossils: common neochiastozygus perfectus and prinsius martinii (nannofossils), globanomalina ehrenbergi, bulimina trigonalis (common) and gavelinella danica (foraminifera) with reworked danian and upper cretaceous nannoand microfossils (danian limestone and tor formation equivalent) and macrofossil debris. æbelø formation (middle selandian) – the æbelø formation (heilmann-clausen et al. 1985) comprises slightly calcareous to non-calcareous, silty – very silty, grey clay. biostratigraphic marker fossils: chiasmolithus edentulus, chiasmolithus eograndis (nannofossils), isabelidinium? viborgense and palaeoperidinium pyrophorum (dinocysts). holmehus formation (selandian–thanetian) – the holmehus formation (heilmann-clausen et al. 1985) consists of non-calcareous, greenish, brownish and reddish, very finegrained clay with minor sand and phosphatic or sideritic concretions. biostratigraphic marker fossils: transversopontis pulcheroides, zygodiscus adamas, neococcolithes protenus (nannofossils), cyclammina amplectens, (foraminifera), alisocysta sp. 1 heilmann-clausen 1985 and cerodinium medcalfii (dinocysts). ølst formation (lower ypresian) – the ølst formation (heilmann-clausen et al. 1985) is a grey to almost black, sandy, silty and mainly non-calcareous clay, interbedded with layers and laminae of black or grey volcanic ash. biostratigraphic marker fossils: sphenolithus editus, tribrachiatus orthostylus, lophodolithus nascens (nannofossils), fenestrella antiqua, coscinodiscus morsianus moelleri, triceratium spp. (diatoms), apectodinium hyperacanthum and deflandrea oebisfeldensis (dinocysts). røsnæs clay formation (middle ypresian) – the røsnæs clay formation (dinesen et al. 1977; heilmann-clausen et al. 1985) comprises red, brown and yellow-brown, very finegrained, plastic marine clay with glauconitic beds at the base and greenish volcanic ash layers throughout. fig. 2. chrono-, bioand lithostratigraphy of the campanian–eocene in denmark. note, not all formations have been found in the investigated area. np14 np15 (pars) np10 49.0 e o ce n e np13 np12 np9 np11 np8 np6 np5 np4 np3 np2 np1 np7 nnte 8 (pars) nnte 7 nnte 1 nnte 2 nntp 4 nntp 3 nntp 1 uc16–20 nntp 2 nntp 9 nntp 10 n o t z o n e d 8 7 5 6 6 5 4 3 lillebælt clay fm l2 fur fm ølst fm stolle klint clay østerrende clay holmehus fm æbelø fm danian limestone tor fm equiv. denmark onshore kerteminde marl lellinge greensandp al eo ce n e 60 55 65 65.0 ypresian l o w er m id d le u p p er l o w er 55.5 60.0 selandian 57.9 54.5 c re t. u . 50 lutetian sparnacian thanetian maastrichtian danian campanian/ røsnæs clay fm r5/r6 d6a / d6b upper d9 lower d9 d8 d7a d7b d5b d5a d4 viborg zone 7 viborg zone 6 viborg zone 5 viborg zone 4 viborg zone 3 viborg zone 2 knudsh. mb. røsnæs clay fm r5/r4 l1 r6 r1 h ei lm an n c la u se n ( 1 9 8 5 ) h ei lm an n c la u se n & c o st a (1 9 8 9 ) martini (1971) north sea biozones standard biozones chronostratigraphy lithostrat. danish biozones calcareous nannofossils varol (1998) burnett (1998) dinoflagellate cysts berggren et al. (1995) stageseries a ge ( m a) calcareous nannofossils 15 biostratigraphic marker fossils: discoaster kuepperi, imperiaster obscurus, toweius occulatus, rhabdosphaera pinguis, discoaster lodoensis (nannofossils), turrilina brevispira, clavulina anglica, pseudohastigerina wilcoxensis, gaudryina hiltermanni, subbotina patagonica (foraminifera), dracodinium? condylos, ochetodinium romanum, wetzeliella astra, dracodinium solidum, eatonicysta ursulae and dracodinium varielongitudum (dinocysts). lillebælt clay formation (upper ypresian – lower lutetian) – the lillebælt clay formation is a green-grey, non-calcareous, very fine-grained plastic clay containing concretions, with red-brown clay beds in the lower part (dinesen et al. 1977; heilmann-clausen et al. 1985). biostratigraphic marker fossils: vaginulinopsis decoratus (foraminifera), membranilarnacia compressa, charlesdowniea columna and eatonicysta furensis (dinocysts). geological development lolland is located to the south of the ringkøbing–fyn high, in the german basin. during the late cretaceaous, fully marine conditions characterised the danish area including the northern highs and chalk of the campanian – upper maastrichtian tor formation equivalent was deposited. danian limestone and lellinge greensand have recently been found in situ on southern lolland and are documented here for the first time. the danian limestone is c. 2 m thick. danian nannofosssils and microfossils are also found in small limestone clasts within the selandian deposits, indicating that danian sediments were eroded and partially redeposited at the end of the danian and probably also in the earliest selandian. the lower selandian glauconitic and conglomeratic deposits of the lellinge greensand rest on danian sediments on sjælland, especially in the copenhagen area (stouge et al. 2000). clasts, microfossils and nannofossils from the lower selandian also indicate deposits from the lellinge greensand in the femern bælt area. the lellinge greensand is at least 12–13 m thick; the top is eroded and covered by quaternary deposits. in dgu core 241.213 the lellinge greensand is found 138 m below ground surface. the encountered glauconitic deposits from the lelllinge greensand rest on danian limestone in a small subsided fault block, demonstrating that the lowermost selandian was probably deposited in the area and subsequently eroded. the early–middle paleocene sea was more restricted than in the cretaceous and several highs bordered the marine area (clausen & huuse 2002). the danian was dominated by erosion, rising sea level characterised the early selandian and erosion dominated again during the late early selandian. these two erosional episodes may have resulted in sediment removal, but it is also possible that the area was subaerial during part of the danian–selandian. continued erosion or nondeposition characterised the upper selandian and therefore deposits from the kerteminde marl formation (up to 100 m thick on sjælland) are not found in the area, probably due to uplift during this time. clastic marine sedimentation resumed during the upper selandian sea-level rise, and the north sea, denmark a b c d e f g h i j k l m n o p q r s t u v w x y z æ ø å 20 μm 20 μm 20 μm 20 μm 20 μm 20 μm 20 μm 20 μm20 μm 20 μm 0.2 mm 0.5 mm 5 μm 5 μm 0.2 mm 0.2 mm 0.2 mm0.2 mm 0.2 mm 0.2 mm 5 μm 0.2 mm 10 μm 5 μm 5 μm5 μm 5 μm 0.2 mm 0.2 mm fig. 3. selected important nannofossils (a–h), microfossils (i–s) and dinocysts (t–å). a: discoaster kuepperi, b: tribrachiatus orthostylus, c: lophodolithus nascens, d: neococcolithes protenus, e: toweius occulatus, f: imperiaster obscurus, g: eiffellithus turriseiffelii, h: prediscosphaera spinosa, i: subbotina patagonica, j: clavulina anglica, k: gaudryina hiltermanni, l: fenestrella antiqua, m: coscinodiscus morsianus moelleri, n: triceratium spp., o: bolivina incrassata, p: bolivinoides laevigatus, q: reusella szajnochae szajnochae, r: stensioeina pommerana, s: praebulimina levis, t: eatonicysta furensis, u: dracodinium? condylos, v: ochetodinium romanum, w: apectodinium hyperacanthum, x: deflandrea oebisfeldensis, y: cerodinium medcalfii, z: isabelidinium? viborgense, æ: cannosphaeropsis utinensis, ø: xenascus wetzelii, å: palaeoperidinium pyrophorum. 1616 and the german basin formed a partly enclosed shelf area. the remainder of the palaeogene (æbelø, holmehus, ølst, røsnæs clay and lillebælt clay) formations mainly consist of fine-grained, plastic clay but there were many lithological shifts during this time. deposition occurred in a relatively deep marine basin, although proximal to the shore. the changes in depositional conditions were related to changing circulation patterns in the north sea and shifts in clay mineral provenance and periods of regression and non-deposition also occurred (heilmann-clausen et al. 1985). intense volcanic activity caused by the opening of the north atlantic resulted in ash deposition during the late paleocene and eocene. quaternary glacial activity eroded, faulted and folded the palaeogene sediments resulting in their present distribution (fig. 4). the new biostratigraphic studies indicate that the røsnæs clay formation is strongly folded with repeated stratigraphy, indicating movements along faults. the youngest formation in the area, the lillebælt clay formation, occurs only as floes in the glacial sediments. on land, the palaeogene clay seems to be undisturbed towards the north whereas disturbances increase towards the south. conclusions the recent biostratigraphic study has provided important information about the pre-quaternary deposits under the c. 18 km wide femern bælt, imperative to planning the construction of the femern bælt fixed link. the cretaceous– palaeogene tor formation equivalent, danian limestone, lellinge greensand, æbelø, holmehus, ølst, røsnæs and lillebælt formations have been identified in multiple boreholes using multidisciplinary biostratigraphy. this information, coupled with physical rock properties (rambøll arup jv 2011) allows a geological cross-section of the femern bælt to be established and demonstrates the complex nature of the depositional and structural history of the area. in addition, in situ danian limestone has been discovered for the first time, along with lellinge greensand on lolland in dgu core 241.213. the new data from >500 m fully cored boreholes provide an excellent basis for future detailed biostratigraphic, sedimentological and basin analysis of the palaeogene deposits in the area. references berggren, w.a., kent, d.v., swischer iii, c.c. & aubry, m.-p. 1995: a revised cenozoic geochronology and chronostratigraphy. in: berggren, w.a. et al. (eds): geochronology, time scale and global stratigraphic correlation. society for sedimentary geology (sepm) special publication 54, 129–212. burnett, j.a. 1998: upper cretaceous. in: bown, p.r. (ed.): calcareous nannofossil biostratigraphy. british micropalaeontological society series, 132–199. clausen, o.r. & huuse, m. 2002: mid-paleocene palaeogeograhpy of the danish area. bulletin of the geologocal society of denmark 49, 171–186. deegan, c.e. & scull, b.j. 1977: a standard lithostratigraphical nomenclature for the central and northern north sea. the institute of geological sciences report 77/25, 36 pp. dinesen, a., michelsen, o. & lieberkind, k. 1977: a survey of the paleocene and eocene deposits of jylland and fyn. danmarks geologiske undersøgelse serie b1, 15 pp. gry, h. 1935: petrology of the paleocene sedimentary rocks of denmark. danmarks geologiske undersøgelse ii. række 61, 180 pp. heilmann-clausen, c. 1985: dinoflagellate stratigraphy of the uppermost danian to ypresian in the viborg 1 borehole, central jylland, denmark. danmarks geologiske undersøgelse serie a7, 69 pp. heilmann-clausen, c., nielsen, o.b. & gersner, f. 1985: lithostratigraphy and depositional environments in the upper paleocene and eocene of denmark. bulletin of the geological society of denmark 33, 287–323. heilmann-clausen, c. & costa, l.i. 1989: dinoflagellate zonation of the uppermost paleocene? to lower miocene in the würsterheide research well, northwest germany. geologisches jahrbuch a111, 431–521. håkansson, e., & pedersen, s.s. 1992: kort over den danske undergrund. københavn: varv (map sheet). martini, e. 1971: standard tertiary and quaternary calcareous nannoplankton zonation. in: farinacci, a. (ed.): proceedings of the second planktonic conference roma. edizioni tecnoscienza, rome 2, 739–785. rambøll arup jv 2011: summary of geological conditions. geotechnical data report 01.3-002, 53 pp. virum: femern a/s. sheldon, e. & nøhr-hansen, h. 2010: fehmarn belt fixed link pre-quaternary biostratigraphy – a final status report for rambøll arup joint venture. danmarks og grønlands geologiske undersøgelse rapport 2010/134, 53 pp. stouge, s., hjortkjær, b.f., rasmussen, j.a., roncaglia, l. & sheldon, e. 2000: microand nannofossil biostratigraphy across the danian/selandian (paleocene) stage boundary at gemmas allé, copenhagen, denmark. gff 122, 161–162. stockholm: geological society of sweden. varol, o. 1998: palaeogene. in: bown, p.r. (ed.): calcareous nannofossil biostratigraphy. british micropalaeontological society series, 200–224. quaternary deposits fehmarn 0 m 50 100 150 lolland 3 km folded and faulted palaeogene clays palaeogene formations cretaceous fig. 4. simplified geological cross-section across the femern bælt. the folding shown in the palaeogene clay is for illustrative purposes only; the real nature of the disturbance (folding and faulting) is much more intense and complicated than shown (modified after rambøll arup jv 2011). authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: es@geus.dk geological survey of denmark and greenland bulletin 42, 2018, 15-37 15 biostratigraphy of the hareelv formation (upper jurassic) in the blokelv-1 core, jameson land, central east greenland peter alsen and stefan piasecki the hareelv formation in the blokelv-1 core is biostratigraphically subdivided by means of ammonite and dinoflagellate cyst stratigraphy. the succession ranges from the oxfordian c. densiplicatum chronozone to the volgian p. elegans chronozone. the mudstones of the blokelv-1 core are characterised by large amounts of amorphous organic matter. this hampers the preparation and identification of dinoflagellate cysts, which are also commonly degraded and corroded. ammonites, on the other hand, are common and well-preserved in the core, contrasting with that observed in the equivalent facies and stratigraphic interval at outcrop. integration of the ammonite and dinoflagellate cyst biostratigraphical data yields a robust chronostratigraphic subdivision of the middle oxfordian – lowermost volgian cored section. keywords: jameson land basin, east greenland, ammonites, dinoflagellate cysts, biostratigraphy, chronostratigraphy, oxfordian, kimmeridgian, volgian ___________________________________________________________________________ p.a. & s.p., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: pal@geus.dk s.p., also geological museum, natural history museum of denmark, university of copenhagen, øster voldgade 5–7, dk-1350 copenhagen k, denmark. e-mail: stefan.piasecki@snm.ku.dk ammonites and dinoflagellate cysts have been applied as the principal means of biostratigraphic dating of the fully cored blokelv-1 borehole (geus 511101), which drilled through 234.80 m of the upper jurassic in southern jameson land (figs1–3). it was the first of three core wells drilled by geus in 2008–2010 as part of a campaign that aimed to document the petroleum potential of the upper jurassic mudstone successions in central east greenland and north-east greenland (bojesen-koefoed et al. 2009, 2014). based on outcrop data, these sediments have previously only shown limited source-rock potential (see bojesen-koefoed et al. 2018, this volume), yet they are time-equivalent with the prolific source rocks of the uk and netherlands north sea (kimmeridge clay formation), the norwegian north sea (draupne formation) and the norwegian sea (spekk formation; brekke et al. 1999, 2001; references in bojesen-koefoed et al. 2018, this volume). the present study of this potential source-rock succession in east greenland has its background in the growing industry interest in areas of offshore northeast greenland. the united states geological survey (usgs) assessment of undiscovered oil and gas resources in the arctic promoted the basins of offshore north-east greenland as being amongst the most important frontier areas for petroleum exploration (gautier 2007; gautier et al. 2011). the danmarkshavn and thetis basins are interpreted to include several kilometre-thick mesozoic successions and have been the subject of particular interest (hamann et al. 2005), yet the validity of the source rock represents a key risk factor in these basins. the purpose of the drilling campaign onshore eastern greenland © geus, 2018. geological survey of denmark and greenland bulletin 42, 15–37. available at: www.geus.dk/bulletin42 mailto:pal@geus.dk mailto:stefan.piasecki@snm.ku.dk http://www.geus.dk/bulletin42 1616 therefore was to investigate fully the source-rock potential of the upper jurassic succession. a suite of analyses was undertaken in the blokelv-1 borehole study, as presented in this bulletin, in order to characterise this important upper jurassic reference section for the northern north atlantic region; an integrated study of the palynoand ammonite biostratigraphy was a central element in this study and forms the subject of this paper. complementary core analysis programmes were subsequently undertaken by geus in fully cored boreholes drilled in the bernbjerg formation in the wollaston forland basin of north-east greenland to extend knowledge of the jurassic – lower cretaceous petroleum potential both stratigraphically and geographically, and in different basin settings (fig. 1). the results of these subsequent investigations will be presented elsewhere. previous biostratigraphic studies of the upper jurassic in the jameson land basin a very well-preserved upper jurassic ammonite succession has been known from milne land at the western margin of the jameson land basin since the work of spath (1935, 1936). the collections on which this work was based were made by a. rosenkrantz in the scoresby sund area during lauge koch’s 1926–1927 mapping expedition in eastern greenland (rosenkrantz 1929). spath’s two monographs describe the ammonites from the oxfordian – lower kimmeridgian and the upper kimmeridgian – volgian. subsequently, sykes & surlyk (1976) and sykes & callomon (1979) revised the boreal oxfordian ammonite zonation and applied it to the successions in east and north-east greenland. callomon & birkelund (1982) and birkelund & callomon (1985) refined spath’s ammonite stratigraphy in milne land, after revisiting and undertaking bed-by-bed collection in the key sections on the eastern flank of hartz fjeld, and adding material from the visdal, bay fjelde and aldinger elv areas (fig. 2). the resulting stratigraphic scheme remains very robust with an established ammonite faunal succession of 11 oxfordian faunal horizons (faunas 3–13 in callomon & birkelund 1980), 10 kimmeridgian faunal horizons (faunas 14–23 in callomon & birkelund 1980; birkelund et al. 1984; birkelund & callomon 1985) and 24 lower and middle volgian faunal horizons (faunas 24–47 in callomon & birkelund 1982). including two middle jurassic faunas, the succession totals 47 jurassic faunal horizons in milne land. the prefix m is used here to denote the bathonian–volgian milne land faunal horizons (i.e. m-1 to m-47) to clearly differentiate them from the stratigraphically slightly overlapping bajocian– oxfordian j-prefixed ammonite faunal horizons ( j-1 to j-41), described from nearby jameson land (callomon 1993, 2003). this usage follows larsen et al. (2003) and callomon et al. (2015). the upper jurassic ammonite zonation established in milne land offers a robust stratigraphic reference/framework to which studies carried out on the upper jurassic elsewhere in east and northeast greenland can be referred. the late jurassic ammonite fauna in jameson land, where blokelv-1 was drilled, fig. 1. simplified geological map of eastern greenland showing the distribution of mesozoic rocks and major structures. c c' b' a a' 18°w 16°w 74°n 76°n 72°n 22°w 20°w ll e cretaceous jurassic triassic permian pre-permian fault buried deepseated faults stauning alper fault post-devonian main fault liverpool land escarpment saf pdmf lle jameson land liverpool landhurry inletscoresby sund hall bredning carls bjer g f jord kong oscar fjord 100 km 26°w 22°w 24°w fig 1_pal milne land traill ø geographical society ø hold with hope clavering ø wollaston forland pd m f sa f 17 fig. 2. geological map of jameson land showing the location of the blokelv-1 drill site. only selected (named) rivers are indicated. ; ; ; ; ; ; bay fjelde blokel v savoia halvø hal l bredn ing aldinger elv visdal hartz fjeld mudderbugten jameson land h urry in le t liverpool land milne land blokelv-1 major dikes/sills faults ice rivers 22°w 72°n 71°n 72°n 71°30'n 70°30'n 23°w24°w25°w 22°w 25°w quaternary cenozoic paleogene basalts hartz fjeld fm hesteelv fm raukelv fm hareelv fm, salix dal mb hareelv fm, sjællandselv mb hareelv fm, katedralen mb olympen fm fossilbjerget fm/pelion fm niell klinter gp kap stewart gp triassic permian carbonifereous devonian basement 20 km kap leslie fm charcot bugt fm paleogene core-well pal fig 2 71°30'nmajor dyke/sill fault ice river (selected) quaternary hesteelv fm hartz fjeld fm palaeogene basalts palaeogene cenozoic raukelv fm hareelv fm, salix dal mb olympen fm fossilbjerget fm/pelion fm neill klinter gp kap stewart gp triassic charcot bugt fm kap leslie fm basement hareelv fm, sjællandselv mb hareelv fm, katedralen mb devonian carbonifereous permian 20 km s c o r e s b y s u n d 1818 is commonly represented by crushed, incomplete and poorly preserved specimens in contrast to the common 3d-preserved mould and steinkern-preservation of the milne land material. subsequent to the establishment of the ammonite succession and stratigraphy in milne land and jameson land, dinoflagellate stratigraphy was established as a biostratigraphic discipline, and studies of the dinoflagellate stratigraphy of the middle and upper jurassic were undertaken in east greenland by piasecki (1981, 1996), poulsen (1985), smelror (1988), milner & piasecki (1996) and larsen et al. (2003). geological setting the jurassic of jameson land forms part of a several kilometre-thick sedimentary succession in the jameson land basin. a short outline of the geological setting is presented here; a more thorough description is presented in bjerager et al. (2018, this volume). the jameson land basin is bounded to the east by the uplifted basement block of the liverpool land high and to the west along faults towards the caledonian basement west of hall bredning (figs 1,2); the basin is c. 150 km wide. its northern boundary is less well constrained, but is probably situated in the transition area between jameson land and traill ø (dam et al. 1995). to the south, beyond scoresby sund, the basin probably continues beneath the thick palaeogene basalt cover, indicating the basin to be more than 200 km long in an n–s direction (figs 1,2). the basin is tilted, with strata dipping slightly towards the south-west. upper palaeozoic strata are thus exposed in the northern areas and progressively younger strata are preserved and exposed towards the south, so that the youngest unit, a lowermost cretaceous incised valley fill (hesteelv formation), is exposed in a small area in southernmost jameson land (fig. 2). however, valanginian and hauterivian–barremian sedimentary rocks are exposed along the western basin margin of milne land (piasecki 1979; birkelund et al. 1984) and thick sedimentary successions have been recognised in geophysical data below the fjord of scoresby sund (larsen & marcussen 1992). it is thus suggested that cretaceous sedimentation continued in the southern part of the jameson land basin maybe until the early cenozoic; paleocene sediments are recorded to underlie palaeogene flood basalts south of jameson land (nøhr-hansen & piasecki 2002). the location of the blokelv-1 drill site was chosen with the objective of studying the inferred fig 3_pal lithostrat. h ar ee lv fo rm at io n ka te dr ale n m em be r sjæ lla nd se lv m b 10 depth (m) 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 210 220 230 233.8 be be mudstone heterolith (mudstone/sandstone) sandstone sandstone, remobilised (intruded) lithology igneous intrusion be ripple cross-lamination wavy bedding slump sandstone intrusion small mudstone intraclasts bentonite large mudstone clast coalified wood belemnite ammonite bivalve brachiopod structures, biota parallel lamination/ bedding diffuse stratification clay si sand pebbl. fig. 3. lithological log of the hareelv formation in the blokelv-1 core; the succession is dominated by the katedralen member, capped by a thin interval of the sjællandselv member. 19 optimum source-rock interval, i.e. the thickest and most organic-rich mudstone section, which was expected to be found in the deepest, axial part of the basin. the site location at blokelv guaranteed the presence of the katedralen member, and by starting the borehole near the base of the sjællands elv member, it was ensured that the borehole would include the uppermost part of the katedralen member which had been poorly sampled in previous drilling campaigns (see bjerager et al. 2018, this volume). material and methods core material is housed at geus. figured ammonites are assigned with mguh numbers (31799–31821) and are stored in the collections of the geological museum of copenhagen (natural history museum of denmark). ammonites the poor exposures of the hareelv formation around the drill site at blokelv are characterised by a lack of ammonites. during the 2008 drilling operation, attempts to collect and sample ammonites in the vicinity of the drill site were unsuccessful. during drilling, a few ammonites were noted in the core. in the laboratory, the core was searched systematically revealing a large number of ammonite-bearing intervals. the uppermost ammonite occurrence in the core is at a stratigraphic level well below that of the rocks exposed in the surrounding area, which might explain the unsuccessful search for ammonites at outcrop. the mudstones in the uppermost part of the core are less consolidated than in the lower levels of the core, however, and therefore the ammonite preservation potential is markedly lower. the lack of ammonites detected in the upper part of the core thus does not necessarily reflect an interval that was initially barren of ammonites. ammonites in the mudstones of the blokelv core are flattened impressions with the shell material dissolved; the impressions are generally well-preserved, sometimes excellently preserved. fragments of ammonites are common. most ammonites are cut by the drill and therefore are not complete. however, a relatively large number of small complete ammonites are present. the laminated mudstones split naturally along bedding-planes and particularly well along planes where ammonite impressions weaken the bedding. bedding surfaces that were split open during coring or core management, were systematically inspected for macrofossils. in continuous, unbroken core sections, ammonites are sometimes visible in cross-section, since the relief of the ribbing forms a characteristic crenulated pattern; such core pieces were carefully split with hammer and chisel for inspection (fig. 4). a total of 42 levels with ammonites were recorded in the interval between core depths 213.84 m and 81.93 m – i.e. exclusively in the katedralen member of the hareelv formation. additional levels with ammonite fragments were observed, but were not considered of biostratigraphic importance due to poor preservation, and were not sampled. ammonites in the blokelv-1 core have been mostly identified by comparison with key taxonomic literature on east greenland ammonites (spath 1935, 1936; sykes & surlyk 1976; sykes & callomon 1979; callomon 1985). the ammonite zonation for the oxfordian– volgian interval (fig. 5) was established by surlyk et al. fig. 4. core piece viewed from the side showing a cross-section of a strongly ribbed ammonite forming a characteristic crenulated bedding plane (arrow). core diameter: 56 mm. pal fig 4 2020 u pp er ju ra ss ic o xf or di an lo w er lo w er m id dl e up pe r sy st em se rie s te th ys (s ta nd ar d) bo re al su bs ta ge chronozones chronostratigraphy subchronozones fa un al ho riz on s up pe r ki m m er id gia n ti th on ian vo lgi an age (ma) 150 160 155 stage fig 5_pal pectinatites pectinatus pectinatites hudlestoni pectinatites wheatleyensis pectinatites elegans pectinatites scitulus aulacostephanus autissiodorensis aulacostephanus eudoxus aulacostephanus mutabilis rasenia cymodoce pictonia baylei a. rosenkrantzi amoeboceras regulare amoeboceras serratum cardioceras tenuiserratum cardioceras densiplicatum amoeboceras (prionodoceras) glosense m-26 m-25 m-24 m-23 m-22 m-21 m-20 m-19 m-18 m-17 m-16 m-15 m-29 m-30 m-12, -13 m-11 m-10 m-8 m-5 m-7 m-9 + j-41 m-6 + j-40 m-14 p. paravirgatus p. eastlecottensis a. (amoebites) bauhini a. (prionodoceras) marstonense a. (prionodoceras) serratum a. (amoeboceras) koldeweyense a. glosense a. (a.) ilovaiskii c. (cawtonic.) blakei c. (maltoniceras) maltonense c. (vertebriceras) vertebrale c. (miticard.) tenuiserratum m-27 m-28 fig. 5. upper jurassic (middle oxfordian – lower volgian) ammonite chronozone scheme for east and north-east greenland. the vertical scale follows the geochronology of gradstein et al. (2012). the chronozone breakdown is based on surlyk (1978, 1991), sykes & callomon (1979), callomon & birkelund (1980, 1982), birkelund et al. (1984) and birkelund & callomon (1985). 21 (1973), surlyk (1978), callomon & birkelund (1980, 1982), birkelund et al. (1984), birkelund & callomon (1985) and summarised in surlyk (1991, fig. 6); it was recently reviewed by kelly et al. (2015). dinoflagellate cysts fifty-eight mudstone samples were collected from the core for palynological analysis and processed using standard preparation techniques in the stratigraphic laboratory at geus. the preparation process includes treatment with acids (hcl, hf, hno3) and filtering with 20 µm filters. this treatment removes carbonates and silicates (clay, silt and sand) from the samples and the remaining organic sedimentary material is resistant to the acids. the amorphous organic material from the mudstone samples is highly resistant to this preparation such that repeated oxidation, extended ultrasonic treatment, washing in potassium hydroxide and filtering were necessary to recover at least some identifiable dinoflagellate cysts. the organic residue was mounted in glycerinegelatine on glass microscope slides for visual analysis by light microscopy. due to inadequate break-down and removal of the amorphous organic matter in most samples from the core, the record of dinoflagellate cysts is sporadic and productive samples are randomly distributed up the core section. furthermore, the identification of species is hampered by the high content of amorphous organic matter (fig. 6). the dinoflagellates are also strongly corroded and badly preserved, and thin-walled specimens fig 6_pal e ~50 µm ~50 µm a b c d e f fig. 6. a–d: preservation of earliest kimmeridgian (p. baylei chronozone) fossil dinoflagellate cysts in the blokelv-1 core (b, d; core depth 102.06 m) compared with contemporaneous specimens of the same species from outcrop material from milne land (a, c; kap leslie formation (bays elv member), ggu 245830). a, b: gonyaulacysta jurassica; c, d: taeniophora sp.; the scale-bar in a is applicable to a–d. e, f: two contrasting palynomorph assemblages from the hareelv formation in blokelv-1; the scale-bar in e is also applicable to f. e: assemblage dominated by marine organic matter, particularly amorphous organic material (aom), sample 232.05 m. f: assemblage rich in black woody material, sample 17.00 m. 2222 20 40 60 80 100 120 140 160 180 200 220 td u pp er ju ra ss ic ki m m er id gia n o xf or di an p. elegans a. autissiodorensis a. mutabilis p. baylei – r. cymodoce p. baylei a. regulare – a. rosenkrantzi a. glosense – a. serratum c. tenuiserratum c. densiplicatum 0 volgian ? a. eudoxus r. cymodoce pe rio d/ ep oc h ag e chronozones lithology core samples (metres) ca rd ioc er as (c aw to nic er as ) a ff. bla ke i/c aw to ne ns e am oe bo ce ra s ( pr ion od oc er as ) c f. g los en se am oe bo ce ra s ( pr ion od oc er as ) s er ra tu m am oe bo ce ra s ( pr ion od oc er as ) c f. s er ra tu m am oe bo ce ra s s p. in de t. am oe bo ce ra s a ff. glo se ns e am oe bo ce ra s c f. s er ra tu m /k old ew ey en se am oe bo ce ra s c f. m ar sto ne ns e am oe bo ce ra s r eg ula re ri ng ste ad ia sp . am oe bo ce ra s r os en kr an tz i am oe bo ce ra s c f. f re bo ldi am oe bo ce ra s f re bo ldi am oe bo ce ra s c f. r eg ula re am oe bo ce ra s c f. r os en kr an tz i am oe bo ce ra s ( am oe bit es ) c f. b au hin i am oe bo ce ra s ( am oe bit es ) a ff. er ne sti am oe bo ce ra s ( am oe bit es ) c f. e rn es ti pic to nia c f. s p. a b irk el un d & c all om on 1 98 5 pa ch yp ict on ia? sp . au lac os te ph an iod es c f. m ut ab ilis au lac os te ph an iod es sp . in de t. am oe bo ce ra s ( eu pr ion od oc er as ) c f. k oc hi 81.93 85.58 85.87 86.14 87.63 91.87 97.34 108.01 108.93 109.25 110.05 119.85 119.89 120.29 122.08 123.53 125.43 145.65 147.99 154.34 155.65 157.81 157.83 158.80 159.53 169.17 201.01 205.30 206.08 208.08 208.13 211.30 213.84 ammonites d ep th (m et re s) pal fig.7 cl.si. sa. peb. fig. 7. chart showing the stratigraphic distribution of ammonites in the blokelv-1 core. for lithological legend, see fig. 3. note that focus is on displaying the relative succession of the individual taxa; where occurrences are closely spaced, the expanded sample position shown in the depth column is utilised in the distribution chart. 23 may also be degraded due to heavy oxidation during preparation in the laboratory. despite these setbacks, a number of important stratigraphic events can be recognised, especially when based on robust and abundant species. in view of the preservation and processing problems outlined above, semi-quantitative palynological analysis was considered to lack statistical significance and thus was not employed. in northwest europe, jurassic dinoflagellate biostratigraphy has been correlated to a subboreal–boreal ammonite zonation (e.g. woollam & riding 1983; riding & thomas 1992). however, the northward continuation of the jurassic seaway (ziegler 1982) from warm tethyan to boreal and possible arctic environments resulted in stratigraphic variations in the first and last occurrences of the dinoflagellate species at different latitudes. the ammonite stratigraphy in east greenland clearly illustrates the interaction of subboreal and boreal faunas (callomon 1993; callomon & birkelund 1982), and the increasing boreal affinities towards the north are clearly demonstrated in the ammonite and dinoflagellate assemblages recorded from northern norway (wierzbowski et al. 2002), the wandel sea basin, north greenland (håkansson et al. 1981) and svalbard (århus 1988). other fossil groups in addition to ammonites, macrofossils in the core include belemnites, represented by rostra (few) and onychites (arm hooks from belemnite cephalopods), bivalves and rare vertebrate remains (bone fragments and teeth). mudstones were sampled for microfossil and nannofossil analysis at three levels (core depths: 206 m, 198 m, 190 m), but neither foraminifers nor coccoliths were recovered (e. sheldon, personal communication 2009). the apparent absence of calcareous microscopic fossils in the blokelv-1 core is probably a function of dissolution as a result of the high content of pyrite and the consequent acidic conditions in the sediment. stratigraphic methods boreal vs tethyan stage nomenclature due to marked faunal provincialism, the uppermost jurassic – lowermost cretaceous interval is commonly subdivided differently in the tethyan and boreal realms (fig. 4; see discussion in zeiss 2003; surlyk 2003). this presently unresolved and controversial subject will not be considered further here; the boreal subdivision (kimmeridgian–volgian–ryazanian) has been routinely applied in east and north-east greenland, and is adopted in this study and parallel studies in north-east greenland. ammonite chronozones an additional area of debate amongst ammonite palaeontologists/stratigraphers in particular concerns the definitions of, and relationships between, chronozones, standard zones and biozones (see callomon 2003; page 2003; zeiss 2003). certain schools consider the jurassic ammonite stratigraphic record to be so highly resolved and well-studied as to define discrete chronostratigraphic units, variably termed chronozones or standard zones. such rock units, in principle, represent a period of time that can also be recognised using fossil groups other than ammonites. this stratigraphic concept has been widely applied in east greenland (piasecki et al. 2004a, 2004b; piasecki & stemmerik 2004; vosgerau et al. 2004). in milne land, in particular, the detailed upper jurassic ammonite chronozone classification has been integrated with the dinoflagellate record (e.g. piasecki 1981; larsen et al. 2003). dinoflagellate events and ranges are closely merged with the ammonite faunas and zonation, the former commonly being based on sample material derived directly from ammonite specimens (e.g. larsen et al. 2003; piasecki et al. 2004b). rather than developing a local palynological biozonation for milne land, the dinoflagellate data were related directly to the ammonite chronozones (piasecki 1981; larsen et al. 2003). this approach is maintained in this study; given the reference chrono-zonation of the milne land succession, the ammonite and palynological dataset presented here permits the breakdown of the blokelv core section into ammonite chronozones. it should be noted that a chronozone may be recognized based on either key ammonites or dinoflagellates, or on a combination of both groups. integrated ammonite and dino flagellate stratigraphy the biostratigraphic subdivision of the blokelv-1 core is described from total depth (td) at 233.80 m and upwards. the recorded ammonite levels are listed in table 1 and illustrated in a stratigraphic distribution chart (fig. 7). the dinoflagellate records and events are illustrated in a stratigraphic distribution chart (fig. 8). 2424 c. densiplicatum chronozone (233.80 (td) – 217.00 m) ammonites were not recorded in this chronozone and its recognition is based on its dinoflagellate record. the lower boundary is placed arbitrarily at the base of the cored section (td), exactly 1 m below the lowermost sample containing palynomorphs (232.80 m). dinoflagellates assemblage. a poor assemblage characterised by the highest occurrences of species that generally show last occurrences in the oxfordian: kalyptea spp., nannoceratopsis pellucida, rigaudella aemula, trichodinium scarburghense and wanaea spp. stratigraphy. the last occurrence of trichodinium scarburghense (232.80 m), rigaudella aemula (224.80 m) and the maximum occurrence of kalyptea spp. (218.40 m) are all events that were reported in milne land in the upper c. densiplicatum chronozone between ammonite faunal horizons m-6 and m-7 (piasecki 1996). age middle oxfordian, late jurassic. c. tenuiserratum chronozone (217.00–208.13 m) the lower boundary of the chronozone at 217 m is placed arbitrarily between the highest occurrence of the kalyptea spp. acme event at 218.40 m (in the under-lying c. densiplicatum chronozone) and the lowest occurence of the ammonite c. (c.) aff. blakei spath 1935 or cawtonense (blake & huddleston 1877) at 213.84 m. depth (m) table 1. core depth and stratigraphy of ammonites in blokelv-1 ammonite taxon figure chronozonefaunal horizon mguh no. 81.93 amoeboceras (euprionodoceras) cf. kochi spath fig. 10i m-20 a. eudoxus 85.58 aulacostephanoides cf. mutabilis (sowerby) m-19 a. mutabilis 85.87 aulacostephanoides cf. mutabilis (sowerby) fig. 10h m-19 86.14 aulacostephanoides sp. indet. 87.63 aulacostephaniodes cf. mutabilis (sowerby) fig. 10g m-19 a. mutabilis 91.87 pachypictonia? sp. fig. 10f m-16 r. cymodoce 97.34 amoeboceras sp. indet. 108.01 amoeboceras sp. indet. 108.93 amoeboceras sp. indet. 109.25 amoeboceras (amoebites) cf. ernsti (fischer) fig. 10e m-14 p. baylei 110.05 pictonia cf. sp. a. birkelund & callomon 1985 fig. 10d m-14 p. baylei 119.85 amoeboceras (amoebites) cf. ernesti (fischer) fig. 10c m-14 p. baylei 119.89 amoeboceras (amoebites) aff. ernesti (fischer) fig. 10b m-14 p. baylei 120.29 amoeboceras (amoebites) aff. ernesti (fischer) fig. 10a m-14 p. baylei 122.08 amoeboceras (amoebites) cf. bauhini (oppel) m-12, -13 a. regulare – a. rosenkrantzi 123.53 amoeboceras cf. rosenkrantzi spath fig. 9l m-12, -13 a. regulare – a. rosenkrantzi 125.43 amoeboceras cf. regulare spath fig. 9k m-12, -13 a. regulare – a. rosenkrantzi 145.65 amoeboceras rosenkrantzi spath fig. 9j m-12, -13 a. regulare – a. rosenkrantzi 147.99 amoeboceras freboldi spath fig. 9i m-12, -13 a. regulare – a. rosenkrantzi 154.34 amoeboceras cf. freboldi spath m-12, -13 a. regulare – a. rosenkrantzi 155.65 amoeboceras rosenkrantzi spath fig. 9h m-12, -13 a. regulare – a. rosenkrantzi 157.81 ringsteadia sp. m-12, -13 a. regulare – a. rosenkrantzi 157.83 ringsteadia sp. m-12, -13 a. regulare – a. rosenkrantzi 158.80 amoeboceras regulare spath fig. 9g m-12, -13 a. regulare – a. rosenkrantzi 159.53 amoeboceras cf. marstonense spath fig. 9f m-12, -13 a. regulare – a. rosenkrantzi 169.17 amoeboceras cf. serratum (sowerby) or koldeweyense sykes & callomon fig. 9e m-11 a. glosense – a. serratum 201.01 amoeboceras aff. glosense (bigot & brasil) fig. 9d m-11 a. glosense – a. serratum 205.30 amoeboceras sp. indet. 206.08 amoeboceras (prionodoceras) cf. serratum (sowerby) fig. 9c m-11 a. glosense – a. serratum 208.08 amoeboceras (prionodoceras) serratum (sowerby) fig. 9b m-11 a. glosense – a. serratum 208.13 amoeboceras (prionodoceras) cf. glosense (bigot & brasil) fig. 9a m-10, -11 a. glosense – a. serratum 211.30 cardioceras (cawtoniceras) aff. blakei spath or cawtonense (blake & huddleston) fig. 8c, d m-8 c. tenuiserratum 213.84 cardioceras (cawtoniceras) aff. blakei spath or cawtonense (blake & huddleston) fig. 8a, b m-8 c. tenuiserratum 31821 31820 31819 31818 31817 31816 31815 31814 31813 31812 31811 31810 31809 31808 31807 31806 31805 31804 31803 31802 31801 31800 31799 pal table 1 25 ammonites an ammonite specimen at 213.84 m is a well-preserved and complete specimen with a relatively narrow umbilicus. it has 17 primary ribs, which on the last whorl develop from fine, slightly bullate to markedly thick, slightly sinuous, bullate ribs. the primaries bifurcate high on the sides and with intercalatories result in 45–50 secondaries. (figs 9a, b). it resembles c. (c.) blakei in ribbing density and ribbing sinuousity but is ribbed in the umbilicus and thus differs from the smooth umbilicus in a specimen figured by sykes & callomon (1979, plate 113, fig. 3). it also resembles c. (cawtoniceras) cawtonense (e.g. callomon 1985, text fig. 8u) which has a ribbed umbilicus, but with straight ribs. an ammonite specimen sampled slightly higher, at 211.30 m, is a small juvenile with relatively weak ribbing on the sides becoming stronger towards the ventrolateral margin (figs 9c, d). the ribs are sinuous and resemble c. blakei. the size difference, however, does not allow direct comparison. both specimens are referred to cardioceras (cawtoniceras) aff. blakei or c. (c.) cawtonense and indicate the faunal horizon m-8, which can be referred to a level in the upper middle oxfordian c. tenuisserratum chronozone (fig. 5). dinoflagellates assemblage. the recorded dinoflagellate assemblage is poor, including kalyptea spp., pareodinia spp. (e.g. p. prologata) and rhyncodiniopsis cladophora. stratigraphy. the c. tenuiserratum chronozone is indicated by the last occurrence of pareodinia prologata at 209.80 m (riding & thomas 1992). age middle oxfordian, late jurassic. a. glosense – a. serratum chronozones (208.13–159.53 m) ammonite occurrences are restricted to the lower part of this interval and the dinoflagellate cyst assemblages are generally poor. the interval is thus referred undifferentiated to the a. glosense – a. serratum chronozones. the lower boundary is placed at the lowest occurrence of a. (p.) cf. glosense at 208.13 m. ammonites a well-preserved, large (size of complete specimen estimated at 82 mm) ammonite at a depth of 208.13 m has rursiradiate ribs developed on the umbilical wall, which curve when crossing the umbilical shoulder and become straight and rectiradiate on the sides (fig. 10a). the primary ribs occasionally bifurcate in the upper part of the flank on the second last whorl. fifteen primaries can be counted on half a whorl, suggesting 30 primaries per whorl. ribbing density increases in the last part of the body chamber, with 10 primaries counted on a quarter of a whorl. the ribs bend strongly forward when crossing the ventral shoulder. the keel appears high. it represents amoeboceras (prionodoceras) cf. glosense (bigot & brasil 1904), indicating the faunal horizons m-10 or m-11, since this species is known to occur in both horizons. these two faunal horizons represent the a. serratum chronozone and the uppermost part of the underlying a. glosense chronozone in greenland. the specimen probably indicates the lower part of that interval since ammonites in the overlying interval, up to 169.67 m, include specimens indicative of the faunal horizon m-11. a. (prionodoceras) serratum (sowerby 1813) is large, with dense, straight and strong prorsiradiate ribs with umbilical, middle and ventrolateral tubercles on the inner whorls, and primaries that bifurcate just below the ventral shoulder (208.08 m, fig. 10b). the ribbing is less strong in the outer whorls, leaving only faint lirae or growth lines. a specimen at 206.08 m that is also characterised by lirae and sometimes by flared ribs is referred to a. (p.) cf. serratum (fig. 10c). a well-preserved microconch with lappet at 201.01 m (fig. 10d) has dense, fine ribbing with backwards-curving primaries on the umbilical shoulder, then becoming straight or gently concave on the flanks, then projected high on the flank and when crossing the ventral shoulder. secondaries appear high on the sides. the ribbing resembles that of a densely ribbed variety of a. glosense figured in sykes & callomon (1979, plate 116, fig. 2) but differs in being much larger than adult microconchs of that species. it is thus referred here to amoeboceras aff. glosense. the highest ammonite indicative of the a. glosense – a. serratum chronozone is found at a depth of 169.17 m (fig. 10e). the specimen is overprinted by trace fossils, but is otherwise well-preserved. it appears to be rather weakly ornamented, but on the last whorl relatively strong, curved tuberculate ribs on the umbilical shoulder are developed; it is otherwise almost smooth on the sides except for faint lirae or growth lines and relatively strong, well-spaced, bullae on the mid-flank. forward-curving 2626 pal fig 8 20 40 60 80 1 00 1 20 1 40 1 60 1 80 2 00 2 20 td upper jurassicperiod/epoch depth (metres) kimmeridgian oxfordianage c hr on oz on es p. ele ga ns a. au tis sio do re ns is a. m ut ab ilis p. ba yle i – r. cy m od oc e p. ba yle i a. re gu lar e – a. ro se nk ra nt zi a . g los en se – a. se rra tu m c. te nu ise rra tu m c. d en sip lic at um d in of lag el lat e ev en ts ab un da nt o lig os ph ae rid ium pa tu lum ? ba se to p rh yn ch od ini op sis m ar to ne ns e ol igo sp ha er idi um p at ulu m pe ris se ias ph ae rid ium pa nn os um ta en iop ho ra sp . m ax im um le pt od ini um su bt ile sy ste m at op ho ra sp p. ka lyp te a sp . m ax im um ri ga ud ell a ae m ula tr ich od ini um sc ar bu rg he ns e c or e sa m pl es (m et re s) ambonosphaera staffinense chytroeisphaeridia hyalina gonyaulacysta jurassica kalyptea spp. nannoceratopsis pellucida pluriarvalium osmingtonense rigaudella aemula sentusidinium spp. sirmiodinium grossii stephanelytron caytonense trichodinium scarburghense wanaea spp. pareodinia ceratophora pareodinia prolongata endoscrinium galeritum kalyptea stegasta pareodinia spp. kalyptea diceras rhynchodiniopsis cladophora systematophora spp. leptodinium subtile circulodinium spp. evansia spp. pareodinia sp. 1 scriniodinium crystallinum gonyaulacysta spp. senoniasphaera spp. protobatioladinium spp. cribroperidinium spp. taeniophora sp. circulodinium distinctum gonyaulacysta dualis apteodinium spp. glossodinium dimorphum gochteodinia mutabilis perisseiasphaeridium pannosum cribroperidinium granuligerum oligosphaeridium patulum rhynchodiniopsis pennata scriniodinium spp. valensiella spp. rhynchodiniopsis spp. senoniasphaera clavelli trichodinium spp. occisucysta monoheuriskos rhynchodiniopsis martonense cribroperidinium sp. a epiplosphaera spp. occisucysta spp. 0. 0 2. 67 2. 96 10 .2 0 13 .7 1 15 .0 0 18 .0 8 20 .1 5 32 .1 0 35 .0 5 47 .0 8 50 .2 7 59 .3 9 64 .9 1 68 .9 7 86 .5 6 91 .2 5 11 4. 60 12 0. 06 12 7. 56 13 5. 41 14 5. 46 15 3. 29 16 0. 80 16 7. 92 17 4. 46 18 5. 11 19 0. 61 19 5. 26 19 8. 25 20 9. 80 21 8. 40 22 4. 80 23 2. 80 d in of lag el lat e cy st s baltisphaeridium spp. veryhachium spp. micrhystridium spp. acritarch spp. ac rit ar ch s algae spp. algae 0 82 .1 4 vo lgi an ? a. eu do xu s r. cy m od oc e ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? 27 bullae are developed on the ventral shoulder. the keel is low and serrated. it is referred here to amoeboceras cf. serratum or a. koldeweyense (sykes & callomon 1979). dinoflagellates assemblage. a poor assemblage including gonyaulacysta jurassica, pareodinia spp., rhynchodiniopsis cladophora and sirmiodinium grossii. the last occurrences of kalyptea spp., evansia spp. and chytroeisphaeridia hyalina occur in this interval. stratigraphy. in milne land, the first occurrences of leptodinium subtile and systematophora spp. are reported in the upper c. tenuiserratum to lower a. glosense chronozones, between ammonite faunal horizons m-8 and m-9 (piasecki 1996). age late oxfordian, late jurassic. a. regulare – a. rosenkrantzi chronozones (159.53–120.29 m) ammonites are relatively common in the fine-grained parts of the interval, whereas dinoflagellate cysts are rare. the lower boundary is placed at the sole occurrence of the ammonite a. cf. marstonense at 159.53 m. ammonites ammonites of the a. rosenkrantzi – a. regulare chronozones are found between 159.53 m and 123.53 m. the lowermost (159.53 m, fig. 10f) is a small, relatively involute, densely ribbed specimen with backward curving secondaries; this is characteristic of a. marstonense spath 1935. the present specimen is, with some caution, referred to a. cf. marstonense, since its small size makes direct comparison with previously figured specimens difficult. a small part of a well-preserved, relatively large specimen occurs at a depth of 158.80 m (fig. 10g). it has straight, densely and regularly spaced ribs, primaries and intercalatories, which curve strongly forward on the ventral shoulder, but disappear and leave the relatively high a b c d b c fig 9_pal fig. 9. ammonites from the c. tenuiserratum chronozone: cardioceras (cawtoniceras) aff. blakei spath 1935 or cawtonense (blake & huddleston 1877). a, b: mguh 31799 (ex geus 511101433) from level 213.84 m (a: natural size. b: ×2). c, d: mguh 31800 (ex geus 511101-432) from level 211.30 m (c: natural size. d: ×2). facing page: fig. 8. chart showing the stratigraphic distribution of palynomorphs in the blokelv-1 core. note that focus is on displaying the relative succession of the individual taxa; where samples are closely spaced, the expanded sample position shown in the depth column is utilised in the distribution chart. the samples indicated with question marks are those in which identifications are tentative due to poor preservation. 2828 keel with smooth sides, whereas the outer margin of the keel is finely serrated. it belongs to a. regulare spath 1935. another specimen (125.43 m, fig. 10k), small and less well-preserved but apparently also with dense ribbing that is strongly projected at the ventral shoulder, resembles a specimen figured and referred to a. cf. regulare by sykes & callomon (1979, plate 118, fig. 4). ammonite fragments with prorsiradiate concave ribbing with the primaries dividing rather high on the flanks, middle flank or higher, are considered to belong to the genus ringsteadia sp. (157.83 and 157.81 m; not figured). a fragment (c. a quarter of a whorl) of an ammonite is preserved at 155.65 m and a faint imprint of another at a depth of 145.65 m (figs 10h, j, respectively). they are involute, strongly ribbed with fairly coarse, straight to slightly curved ribs on the flanks. weak tubercles may develop at the mid-flank, whereas all ribs become tuberculate at the ventrolateral shoulder, where they curve forward (moderately to strongly) on the venter. they represent a. rosenkrantzi spath 1935. another less wellpreserved specimen at 123.53 m is referred to a. cf. rosenkrantzi (fig. 10l). a medium–large specimen is preserved in part at 147.99 m (fig. 10i). the inner whorls are densely and finely ribbed, whereas the outer whorl is characterised by fairly faint ribbing on the flank, but with strong bullate tubercles developed on the ventrolateral margin. it resembles closely the a. freboldi figured by sykes & surlyk (1976, fig. 5f ), and is referred to that species. another less well-preserved specimen is more cautiously referred to as a. cf. freboldi (154.34 m, not figured). a specimen of amoeboceras cf. bauhini (oppel 1863) is recorded at a depth of 122.08 m (not figured). it is crushed and poorly preserved. the ribbing is strong, interrupted by a smooth band two-thirds up the flank. ribbing becomes strong again above the smooth band and persists to the venter, a characteristic of a. bauhini. the preservation, however, only permits tentative identification. the ammonite material within the interval thus indicates the undifferentiated faunal horizons m-12 to m-13 of callomon & birkelund (1980; fig. 4). in britain, the ammonite a. regulare is known to range only to the top of the a. regulare chronozone (sykes & callomon 1979), suggesting that the boundary between the a. regulare and a. rosenkrantzi chronozones may lie between 125.43 m and 123.53 m in the blokelv-1 core. dinoflagellates assemblage. a poor assemblage of cribroperidinium spp., endoscrinium galeritum, gonyaulacysta jurassica, pareodinia spp., rhynchodiniopsis cladophora and scriniodinium crystallinum. stratigraphy. dinoflagellate cysts are very rare in this interval, combined with a low diversity of species. impoverished assemblages are also recorded on milne land, possibly reflecting increased boreal affinities in this interval (sykes & callomon 1979). similar impoverished assemblages are recorded to the north in peary land (håkansson et al. 1981), nordland, norway (wierzbowski et al. 2002) and svalbard (århus 1988). all species recorded from this interval have long ranges and cannot be referred to ammonite chronozones. the broad interpretation of the age based on palynology is middle to late oxfordian. age late oxfordian, late jurassic. facing page: fig. 10. ammonites from the a. glosense – a. serratum chronozones interval and the a. regulare – a. rosenkrantzi chronozones interval in the blokelv-1 core; all figured at natural size. a: amoeboceras (prionodoceras) cf. glosense (bigot & brasil 1904), mguh 31801 (ex geus 511101-426), core depth 208.13 m. b: amoeboceras (prionodoceras) serratum (sowerby 1813), mguh 31802 (ex geus 511101-425), core depth 208.08 m. c: amoeboceras (prionodoceras) cf. serratum, mguh 31804 (ex geus 511101-430), core depth 206.08 m. d: amoeboceras aff. glosense, mguh 31804 (ex geus 511101-431), core depth 201.01 m. e: a. cf. serratum or koldeweyense sykes & callomon 1979, mguh 31805 (ex geus 511101-424), core depth 169.17 m. f: amoeboceras cf. marstonense spath 1935, mguh 31806 (ex geus 511101-739), core depth 159.53 m. g: amoeboceras regulare spath 1935, mguh 31807 (ex geus 511101-457), core depth 158.80 m. h: amoeboceras rosenkrantzi spath 1935, mguh 31808 (ex geus 511101-435), core depth 155.65 m. i: amoeboceras freboldi spath 1935, mguh 31809 (ex geus 511101-434), core depth 147.99 m. j: amoeboceras rosenkrantzi, mguh 31810 (ex geus 511101-454), core depth 145.65 m. k: amoeboceras cf. regulare, mguh 31811 (ex geus 511101-451), core depth 125.43 m. l: amoeboceras cf. rosenkrantzi, mguh 31812 (ex geus 511101449), core depth 123.53 m. 29 fig 10_pal j k l g h i d e f a b c 3030 a b c fig 11_pal d e f g h i fig. 11. ammonites from the p. baylei chronozone – a. eudoxus chronozone interval in the blokelv-1 core; all figured at natural size. a, b: amoeboceras (amoebites) aff. ernesti (fischer 1913). a: mguh 31813 (ex geus 511101-446), core depth 120.29 m. b: mguh 31814 (ex geus 511101-445), core depth 119.89 m. c, e: amoeboceras (amoebites) cf. ernesti. c: mguh 31815 (ex geus 511101-444), core depth 119.85 m. e: mguh 31817 (ex geus 551101417), core depth 109.25 m. d: pictonia cf. sp. a birkelund & callomon 1985, mguh 31816 (ex geus 511101-412), core depth 110.05 m. f: pachypictonia? sp., mguh 31818 (ex geus 511101-420), core depth 91.87 m. g, h: aulacostephaniodes cf. mutabilis (sowerby 1823). g: mguh 31819 (ex geus 511101-402), core depth 87.63 m. h: mguh 31820 (ex geus 511101-407,) core depth 85.87 m. i: amoeboceras (euprionodoceras) cf. kochi spath1935, mguh 31821 (ex geus 511101-413), core depth 81.93 m. 31 p. baylei chronozone (120.29–109.25 m) the base of the chronozone in the blokelv-1 core is placed at the first stratigraphic occurrence of the ammonite amoeboceras aff. ernesti at 120.29 m. ammonites in greenland, the p. baylei chronozone is represented by one faunal horizon, m-14 (fig. 5). the presence of this horizon in the blokelv-1 core is indicated by the occurrence of small amoeboceras microconchs related to a. (amoebites) ernesti (fischer 1913) at four levels (120.29 m, 119.89 m, 119.85 m, 109.25 m). they are characterised by smooth early whorls and isocostate ribbing (figs 11a–c, e). within this interval characterised by a. ernesti, a faint imprint of a specimen of pictonia was recorded at a depth of 110.05 m (fig. 11d). it is evolute and relatively densely and delicately ribbed. the ribs appear to bifurcate high on the outer whorl leaving the bifurcation level hidden on earlier whorls. the ribs are slightly prorsiradiate and convex. the specimen closely resembles p. cf. sp. a as figured by birkelund & callomon (1985, plate 11, fig. 3). dinoflagellates assemblage. a poor assemblage of common taeniophora sp. (fig. 6d) and rare ambonosphaera staffinensis, circulodinium spp. and gonyaulacysta jurassica (fig. 6b). stratigraphy. in milne land, the maximum abundance of taeniophora sp. was recorded in the lower p. baylei chronozone below ammonite faunal horizon m-14 (piasecki 1996); this event is observed at 120.06 m in the blokelv-1 core section. age earliest kimmeridgian, late jurassic. p. baylei – r. cymodoce chronozones (109.25–91.87 m) the lower boundary of this combined chronozone interval is placed at the highest occurrence of a. cf. ernesti at 109.25 m. the ammonite fauna within the interval is non-diagnostic, being referred to amoeboceras spp., and palynological samples yielded only rare algae. no diagnostic biostratigraphic events were recorded, and the interval is defined by the top of the p. baylei chronozone beneath (last occurrence of a. cf. ernesti at 109.25 m) and the base of the r. cymodoce chronozone above (occurrence of pachypictonia? at 91.87 m). r. cymodoce chronozone (91.87– 87.63 m) a relatively thin interval is referred to the r. cymodoce chronozone based on the presence of the ammonite pachypictonia? which defines the base of the chronozone at 91.87 m. ammonites part of a large specimen was recorded at 91.87 m (fig. 11f). it is evolute with strong, almost bullate primary ribs and weak secondaries. the level of furcation on the inner whorls is just hidden by the subsequent whorls, hence the strong primaries dominate the open, shallow umbilicus. the specimen is not sufficiently well preserved to identify it to species level, but its resemblance to ?pachypictonia sp. c. birkelund & callomon 1985 (plate 15, fig. 1) suggests it is assignable to pachypictonia? it is considered to indicate faunal horizon m-16 within the r. cymodoce chronozone. dinoflagellates assemblage. a poor assemblage of glossodinium dimorphum, gonyaulacysta jurassica and sirmiodinium grossii was recorded in this chronozone. age early kimmeridgian, late jurassic. a. mutabilis chronozone (87.63–82.14 m) ammonites are relatively common, occurring at four levels within this thin interval. the base of the a. mutabilis chronozone is placed at the lowest occurrence of the ammonite a. cf. mutabilis at 87.63 m. ammonites ammonites that occur at four levels between 87.63 m and 85.58 m (figs 11g, h) include specimens of aulacostephanus cf. mutabilis (sowerby 1823). the specimens 3232 are parts of large, evolute forms. inner whorls have strong and bullate primaries that curve on the lower flank and bifurcate just below the umbilical seam of the subsequent whorl. later ribbing becomes less strong to weak, and primaries divide into three to four secondaries. in view of the fragmentary preservation, the material is referred to as a. cf. mutabilis. a. mutabilis represents faunal horizon m-19 and the a. mutabilis chronozone. dinoflagellates assemblage. a poor assemblage of gonyaulacysta jurassica, sirmiodinium grossii, endoscrinium galeritum and systematophora spp. was recorded in this chronozone. age middle kimmeridgian, late jurassic a. eudoxus chronozone (82.14–59.39 m) ammonites were recorded at only one level in the a. eudoxus chronozone and the chronozone definition and age assignment rest primarily on the dinoflagellate cyst assemblages. the base of the chronozone is placed at the lowest occurrence of the dinoflagellate p. pannosum at 82.14 m. ammonites the uppermost ammonite-bearing level in the blokelv-1 core is at 81.93 m, just 21 cm above the base of the a. eudoxus chronozone. the ammonite is crushed and fragmented (fig. 11i). ribs are dense and regularly spaced. the degree of involution, ribbing density and ribs that are curved on the umbilical shoulder and then straight and almost rectiradiate on the flanks suggest identification of amoeboceras (euprionodoceras) cf. kochi spath 1935. a. kochi is the index of the lowermost horizon (m20) of the a. eudoxus chronozone. dinoflagellates assemblage. dominated by perisseiasphaeridium pannosum in association with circulodinium distinctum and cribroperidinium spp. stratigraphy. in milne land, the first occurrence of abundant perisseiasphaeridium pannosum is recorded in the lowermost a. eudoxus chronozone, in ammonite faunal horizon m-20 (piasecki 1996) and this event is used in the blokelv-1 section to define the base of the chronozone. the last occurrence of abundant p. pannosum in milne land was recorded near the top of the a. eudoxus chronozone, above ammonite faunal horizon m-22 (piasecki 1996); this event occurs at 68.97 m in the blokelv-1 section. age middle kimmeridgian, late jurassic. a. autissiodorensis chronozone (59.39–10.00) recognition of the chronozone is based on its dinoflagellate cyst assemblages since ammonites were not recorded. the base of the chronozone is placed at the lowest occurrence of the dinoflagellate cyst o. patulum at 59.39 m. dinoflagellates assemblage. dominated by oligosphaeridium patulum and cribroperidinium spp. in association with circulodinium distinctum, perisseiasphaeridium pannosum, rhynchodiniopsis spp. and senoniasphaera clavellii. stratigraphy. in milne land, the first occurrence of abundant oligosphaeridium patulum is recorded at the base of the a. autissiodorensi chronozone between ammonite faunal horizons m-22 and m-23 (piasecki 1996); this event is thus used to place the base of the chronozone at 59.39 m in the blokelv-1 section. age latest kimmeridgian, late jurassic. p. elegans chronozone (10.00–0.00 m) the chronozone boundaries and age of this interval are based on its dinoflagellate cyst assemblages; no ammonites were recorded in this interval. the first (lowest) indication of the chronozone is the appearance of the dinoflagellate cyst r. martonense at 2.96 m, the first occurrence of which is known to be somewhat above the base of the p. elegans chronozone (piasecki 1996). the lower 33 boundary is arbitrarily placed at 10.00 m, at the base of the sjællandselv member. dinoflagellates assemblage. dominated by oligosphaeridium patulum and cribroperidinium spp. in association with circulodinium distinctum, rhynchodiniopsis spp. and senoniasphaera spp. stratigraphy. abundant oligosphaeridium patulum occurs to the top of the blokelv-1 borehole (sample at 0.00 m); note that the top of the recovered cored section is at 1.72 m and the uppermost palynological sample was taken from surface exposure at the drill site. in milne land, the highest occurrence of abundant oligosphaeridium patulum was recorded below the p. wheatleyensis chronozone, below the ammonite faunal horizon m-25 (piasecki 1996). the presence of rhychodiniopsis martonense in the interval is indicative of the p. elegans chronozone as this species was not recorded above the p. elegans chronozone and ammonite faunal horizon m-24 in milne land (piasecki 1996). age earliest volgian, late jurassic. discussion the combined ammonite and dinoflagellate stratigraphy in blokelv-1 provides a detailed subdivision and dating of the drilled succession. separately, ammonite and dinoflagellate stratigraphies would have provided data for only parts of the core. published biostratigraphic studies of the jurassic of jameson land mostly concern macrofossils, especially ammonites. the few published palynological studies have focused particularly on the lower to middle jurassic (e.g. koppelhus & dam 2003; koppelhus & hansen 2003). upper jurassic palynological studies have been presented as survey and consultancy reports and remain unpublished. kelly et al. (2015) recently published a review of the jurassic biostratigraphy of east greenland in which they applied upper jurassic dinoflagellate stratigraphic schemes from two studies related to the north sea/ north atlantic by partington et al. (1993) and poulsen & riding (2003). it is notable, however, that neither of these schemes contain data from the upper jurassic of east greenland. there are undoubtedly many similarities between the dinoflagellate stratigraphies in the north sea/north atlantic and in east greenland, but it is considered somewhat premature to apply distant datasets to another region without the support of local data. stratigraphic conclusions lithostratigraphically, the blokelv-1 borehole encountered two members of the hareelv formation, the kate dralen member (233.80 (td) – 10.00 m) and the sjællandselv member (10.00–0.00 m). the base of the katedralen member was not reached and the unit thus has a thickness in excess of 225 m in this area. this exceeds that predicted for this area based on the first mapping campaign (an estimate of c. 200 m was made by surlyk et al. (1973), but is within the thickness range estimated for the hurry inlet region (200–400 m in surlyk & noenygaard 2001). the cored succession is dated as middle oxfordian to earliest volgian using ammonites and dinoflagellates, and the succession is subdivided into ammonite chronozones. all middle oxfordian – lower volgian chronozones have been identified in the core. the stratigraphic resolution is variable, however, and combined, undifferentiated chronozones were recognised in the upper oxfordian (a. serratum – a. glosense and a. regulare – a. rosenkrantzi) and the lower kimmeridgian (p. baylei – r. cymodoce). given the present stratigraphic resolution, the succession appears to represent continuous deposition from the middle oxfordian to the earliest volgian, as no hiatus was recognised. the zonation allows a detailed correlation of the blokelv-1 core to, and a framework for, other shallow cores and outcrops in the jameson land basin (fig. 12), and it contributes to an understanding of the depositional history of the basin (bjerager et al. 2018, this volume). in addition, it supports correlation of the katedralen member source rock in the blokelv-1 core with ‘kimmeridge clay’ equivalents in the north atlantic region (bojesenkoefoed et al. 2018, this volume) and the barents shelf (leith et al. 1993). the nature of the palynological record supports the interpretation that the sediments of the katedralen member in the blokelv-1 core were deposited in dominantly anoxic bottom conditions that resulted in the preservation of abundant organic material. the oxfordian mudstones contain abundant terrestrial matter compared to the marine-dominated organic matter of the 3434 kimmeridgian mudstones, as confirmed by geochemical analysis (bojesen-koefoed et al. 2018, this volume). nevertheless, the entire mudstone section represents a good– excellent source rock and the difference in source-rock quality between the oxfordian and kimmeridgian mudstones is surprisingly small (bojesen-koefoed et al. 2018, this volume). the sjællandselv member, in contrast, has no potential for hydrocarbons with a low organic content of black woody material. c re ta ce ou s lo w er u pp er ju ra ss ic o xf or di an l l l l m u u u u sy st em se rie s te th ys (s ta nd ar d) ammonite chronozones bo re al su bs ta ge m ki m m er id gia n be rr ias ian ti th on ian ry az an ian vo lgi an age (ma) 140 150 160 145 155 stage fig 12_pal s n raukelv olympen jameson land rasenia cymodoce amoeboceras serratum amoeboceras glosense cardioceras tenuiserratum cardioceras densiplicatum cardioceras cordatum quenstedtoceras mariae pictonia baylei pectinatites elegans pectinatites wheatleyensis pectinatites scitulus pectinatites pectinatus hareelv fm? olympen fm zeus mbzeus mb hareelv fm sa katedralen mb hades mb sjællandselv mb pectinatites hudlestoni a. rosenkrantzi aulacostephanus mutabilis aulacostephanus eudoxus aulacostephanus autissiodorensis amoeboceras regulare shallow marine sandstone shelf transition silty–sandy mudstone, heterolith deep marine sandstone deep shelf (basin) mudstone source rock mass-flow sandstone remobilised/injected sandstone prograding unit raukeelv fm raukeelv fm bl ok el v1 c c hiatus condensed section fig. 12. scheme of the upper jurassic in the jameson land basin showing the stratigraphic position and extent of the blokelv-1 core. 35 acknowledgements jette halskov and stefan sølberg prepared the figures, and annette ryge and dorthe samuelsen prepared the palynological samples. we appreciate the useful comments and suggestions from the two referees w.a.p. wimbledon and m. smelror. references århus, n. 1988. palynostratigraphy of some bathonian–hauterivian sections in the arctic, with emphasis on the janusfjellet formation type section, spitsbergen. iku-report 23.1252.11/01/88, 139 pp. birkelund, t. & callomon, j.h. 1985: the kimmeridgian ammonite faunas of milne land, central east greenland. bulletin grønlands geologiske undersøgelse 153, 56 pp. birkelund, t., callomon, j.h. & fürsich, f.t. 1984: the stratigraphy of the upper jurassic and lower cretaceous sediments of milne land, central east greenland. bulletin grønlands geologiske undersøgelse 147, 56 pp. bjerager, m., alsen, p. bojesen-koefoed, j.a., piasecki, s. & pilgaard, a. 2018: late jurassic evolution of the jameson land basin, east greenland – implications of the blokelv-1 borehole. in: ineson, j. & bojesen-koefoed, j.a. 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(eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 75–114. ziegler, p.a. 1982: geological atlas western and central europe, 130 pp. amsterdam: elsevier. _________________________________________________________________________________________ manuscript received 4 december 2015; revision accepted 1 may 2018 review article | short sandrin 2020: geus bulletin 44. 4836. https://doi.org/10.34194/geusb.v44.4836 1 of 6 semi-conventional play: definition, exploration strategy and the example of the chalk group in denmark alessandro sandrin* department of petrology and economic geology, geological survey of denmark and greenland (geus), copenhagen, denmark abstract play analysis has been widely used in hydrocarbon exploration for decades with great success. in recent years, progress has also been made to describe reservoir properties of very low permeability reservoirs. however, comparatively little research has been conducted into play analysis for such reservoirs, which may lead to misleading estimates of their hydrocarbon potential. here, the concept of a semi-conventional play is defined and characterised as having a reservoir of such low permeability that a hydrocarbon column can form down-dip of an effective dry trap. a new exploration approach is proposed for such plays using the chalk group play in the danish north sea as an example. it is suggested that together with the usual risk elements, a more detailed analysis of ‘charge’ is necessary, paying particular attention to identifying possible hydrocarbon entry-points, palaeostructures and the maximum distance from these entry-points that the hydrocarbons may have reached since they first entered the reservoir. the application of this novel approach for semi-conventional plays in mature basins could help unlock further resources in proximity of existing fields, and reduce the risk of failure in frontier exploration. introduction in petroleum exploration, a ‘play’ is a conceptual model for hydrocarbon accumulations used to identify prospective areas in a basin. the model is further analysed during development of the fields to continue exploitation within a defined geological trend. the hydrocarbon industry developed the concept of play analysis to provide a more structured way of exploring hydrocarbons (white 1988; grant et al. 1996; nelskamp 2017). starting at regional-geology large scales, the analysis moves up within the ‘exploration triangle’ (e.g. fraser 2011; milkov 2015) all the way to prospect analysis and exploratory drilling. individual companies and research centres have developed their own play analysis techniques that differ in details but not in their fundamental aspects. the petroleum play could be divided into a number of play elements, including (1) the reservoir (the rock hosting the hydrocarbons), (2) seal (the cap-rock), (3) trap (a geometrical configuration of rocks that produces a closed volume of hydrocarbons) and (4) charge (how, where and when hydrocarbons moved from the source rock into the reservoir). a key point of the exploration process is the separation between the play elements (regional geological features) and the prospect elements, *correspondence: alsa@geus.dk received: 06 jan 2020 accepted: 23 mar 2020 published: 26 may 2020 keywords: play analysis, chalk, low-permeability reservoir, denmark abbreviations: md: mdarcy twt: two way time geus bulletin is an open access, peerreviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: catherine jex (geus, denmark) reviewed by: ida fabricius (technical university of denmark, denmark) and kresten anderskouv (university of copenhagen, denmark) funding: none declared competing interests: none declared additional files: none provided https://doi.org/10.34194/geusb.v44.4836� mailto:alsa@geus.dk https://creativecommons.org/licenses/by/4.0/deed.ast sandrin 2020: geus bulletin 44. 4836. https://doi.org/10.34194/geusb.v44.4836 2 of 6 www.geusbul let in.org which are specific to a particular area (i.e. in close proximity of the prospect, which is a small and well defined area were hydrocarbons have been predicted to be present). explorationists often assume, but never explicitly state, that the migration of hydrocarbons is basically ‘instantaneous’ on geological timescales. this implies that the reservoir and the hydrocarbon carrier beds must have good permeability. in other words, when a drop of oil enters the reservoir, it is assumed that it instantly reaches the various traps in the play. this is obviously true in reservoirs with permeabilities of the order of 500 md or more (chadwick et al. 2004; blasingame 2008; hjuler et al. 2016). however, in low-permeability reservoirs, hydrocarbons may move at such low rates that dry effective traps could occur and therefore the boundaries of the play have to be re-drawn accordingly. to account for this, a new exploration approach is proposed for semi-conventional plays in this article, using the chalk group in the danish north sea as a type example. the proposed approach may permit a better delineation of the effective play and minimises the chance of leaving economically interesting areas unexplored. definition of semi-conventional play a semi-conventional play is characterised by reservoir permeability so low that the speed at which hydrocarbons move within the reservoir is slow enough to allow the existence of hydrocarbon columns down-flank of a dry effective trap (e.g. see modelling by kok & arnhild 2012). this implies that hydrocarbons can still be ‘on the move’ and may not have yet reached all the available and effective traps. the speed at which hydrocarbons move within the reservoir is not only dependent on permeability but also on other parameters such as pressure gradients, oil viscosity and structural dip (fig. 1; vejbæk et al. 2005; kok & arnhild 2012). however, permeability is a reservoir property that must be accounted for when performing play analysis, and thus it is the focus of this review. the most significant difference between a conventional and a semi-conventional petroleum play is the element of ‘charge’ (table 1). in conventional oil plays, charge normally includes information on source rock presence, quality, maturity, timing and migration. within high-permeability reservoirs, migration is assumed to be ‘instantaneous’ whereas in low-permeability reservoirs migration velocities are of the order of 1–10 km per million years (kok & arnhild 2012). thus, we need to define hydrocarbon entry-points and the distance hydrocarbons may have travelled since the first drop of oil entered the reservoir. a description of charge should thus include a comprehensive analysis of the following: 1. determination of likely hydrocarbon entry-points 2. analysis of pressure gradients driving hydrocarbon migration 3. migration velocity 4. evolution and spatial variation of permeability within the reservoir. these factors are sometimes considered in the analysis of conventional plays. however, they are particularly important when considering semi-conventional plays for reasons described next. hydrocarbon entry-points when performing conventional play analysis, it is often assumed that vertical migration in the reservoir from the source rock happens everywhere within the play area where the source rock is (1) present, (2) of good quality and (3) matured at the right time – in relation to the trap formation. a more robust approach would be to map likely entry-points, which could be, for example, the fractured areas that connect the source rock to the reservoir. for a semi-conventional play, the low migration velocities imply that hydrocarbons may have reached only limited areas within a certain distance from the entry-points. analysis of pressure gradients aquifer pressure gradients can control the movement of hydrocarbons into and within a reservoir (goff 1983; winefield et al. 2005; o’connor et al. 2008). this process is extremely important at a regional scale and determines the preferential direction of hydrocarbon migration. fig. 1 migration velocity of oil for various aquifer pressure gradients and oil saturations. the migration velocity of oil (vo) is calculated using the following equation: vo = 0.21 × (ko/uo) × (dp/dl)/(ø × so), where ko is the oil permeability, uo is the oil viscosity, dp/dl is the pressure gradient, ø is the porosity and so is the oil saturation. https://doi.org/10.34194/geusb.v44.4836 http://www.geusbulletin.org sandrin 2020: geus bulletin 44. 4836. https://doi.org/10.34194/geusb.v44.4836 3 of 6 www.geusbul let in.org moreover, at the fieldor prospect-scale, pressure gradients must be estimated (together with other parameters such as reservoir porosity, tectonic evolution and diagenesis) to characterise the development of oil–water contacts through time and the possible (present-day) occurrence of tilted oil–water contacts (dennis et al. 2000; harris & goldsmith 2001; dennis et al. 2005; vejbæk et al. 2005). migration velocity in low permeability reservoirs, migration velocities could be as low as a few kilometres per million years. mapping of likely entry-points allows for the creation of likelihood maps for the ‘reach’ of hydrocarbons, calculated from various scenarios for migration velocities and directions. here, ‘reach’ is defined as the maximum distance reached by hydrocarbons from likely entry-points. evolution and variation of permeability within the reservoir the reservoir facies may change in space and time. in some areas, the permeability may be lower due to lithological variations. moreover, the permeability may also change in time due to compaction and diagenesis. gross depositional environment maps for the reservoir are required to better predict the permeability of reservoir at various areas. when possible, the reservoir burial history, including compaction and diagenesis, should be analysed by stratigraphic backstripping or structural restorations. the chalk group in denmark the concept of semi-conventional play originated from work on the chalk group play (chalk play) in the danish sector of the north sea (fig. 2). with a purely traditional play analysis approach, major challenges were encountered at both play and prospect scale. the chalk play is a proven play (megson 1992; huuse 1999; frykman et al. 2004; megson & tygesen 2005; van buchem et al. 2017). the first discovery was made in 1966 (kraka field) with the first production at dan field in 1972. the reservoir is a sedimentary carbonate rock, mainly composed of coccoliths (hancock 1975; hardman 1982; fabricius 2007; rasmussen & surlyk 2012) with variable porosity and permeability (see fig. 3 for a comparison between the forties sandstone mbr. and the tor fm. of the chalk group). reservoir layers within the chalk have 30–45% porosity, with 1–10-md permeability (mortensen et al. 1998; fabricius 2007; fabricius et al. 2007). when analysing oil fields in the danish sector, halfdan field (containing c. 1500 million stock barrels; table 1 comparison between play elements for a conventional, unconventional and semi-conventional play conventional petroleum play unconventional petroleum play semi-conventional petroleum play reservoir source rock occurrence reservoir seal hydrocarbon generation seal trap recoverability trap charge   chargea amust include reach (i.e. the maximum distance reached by hydrocarbons from likely entry points), which is dependent on (1) entry points, (2) pressure gradients, (3) migration velocity and (4) permeability variations of the reservoir in time and space. fig. 2 danish fields on top chalk depth map. s, salt structure; csf, coffee soil fault. modified following vejbæk et al. (2005). most of the fields are located on clear structural traps; however, the massive halfdan field is not related to any structural closure and the stratigraphic component of the trapping system cannot fully explain the existence of such a large accumulation. fig. 3 permeability (k) versus porosity for a conventional (forties sandstone mbr.) and a semi-conventional reservoir (tor fm. of the chalk group). note that even for large porosity values (>30%), the tor fm. permeability never exceeds 100 md. data from fabricius (2007) and jones et al. (2005). https://doi.org/10.34194/geusb.v44.4836 http://www.geusbulletin.org sandrin 2020: geus bulletin 44. 4836. https://doi.org/10.34194/geusb.v44.4836 4 of 6 www.geusbul let in.org albrechtsen et al. 2001; fig. 4a) stands out as an evident anomaly since (1) there is no clear structural closure and (2) the stratigraphic component of the trapping system does not fully explain the existence of field. vejbæk et al. (2005) suggested that hydrocarbon column in halfdan field was created before oil reached the up-dip four-way closure of dan field. in their modelling, a palaeoclosure was assumed to exist in halfdan field at the time of charge. however, the tilt or disappearance of the palaeostructure did not result in the ‘instantaneous’ equilibrium of fluids, leaving a substantial hydrocarbon column in the present-day non-existing geometrical trap. this clearly has an impact on both playand prospect-scale analyses of semi-conventional plays in the region. figure 4 depicts a conceptual model for the development of a semi-conventional play through time (figs 4b and c), compared with the present day example of the dan–halfdan fields (fig. 4a). evidently, when performing play analysis, the up-dip four-way structural closure (well-a) may or may not be located within the play (figs 4b and c). in the case of low-permeability reservoirs, hydrocarbon reach could be estimated after mapping a likely entry-point and modelling migration velocities and directions, thus delineating the boundaries of the play ‘sweet spot’ (i.e. the most prospective region). in fig. 4, both well-a and well-b are positioned at the point where the source rock is present and has matured; however, well-b is located in the proximity of a fractured zone that allowed vertical migration of hydrocarbons and the build-up of a column (as modelled by kok & arnhild 2012). at well-a, an effective vertical conduit is missing, which implies that charge can occur only by the lateral movement of hydrocarbons within the reservoir. in a conventional play, a closure, such as at well-a, would be interpreted as being within the play sweet spot, as hydrocarbons entering well-b will reach well-a in short geological times and before the present day. yet, for a semi-conventional play, well-a may be interpreted to be outside the effective play area, fig. 4 a: seismic vertical section in time (twt) across dan and halfdan fields. b: semi-conventional play at the onset of charge, where the oil has built a possible economic accumulation at well-b, but has not yet reached the effective trap at well-a. c: the same semi-conventional play after several million years, where the slow moving oil has reached the effective trap at well-a, and a hydrocarbon column is still present at well-b. see fig. 1 for the location of the profile. [aq10] https://doi.org/10.34194/geusb.v44.4836 http://www.geusbulletin.org sandrin 2020: geus bulletin 44. 4836. https://doi.org/10.34194/geusb.v44.4836 5 of 6 www.geusbul let in.org and thus have a lower chance of success than well-b. this observation is in contrast with the classical approach of prospect risking, since a mapped four-way structural closure above a mature source rock has a lower chance of success than a prospect with no stratigraphic or structural closure. this novel approach to play analysis could be adopted in mature basins to properly evaluate the remaining prospectiveness without overlooking possible economic targets. moreover, this new approach may also be applied in relatively unexplored basins for the correct evaluation of existing wells (i.e. ‘dry well analysis’), for the characterisation of viable prospects and for the delineation of play sweet spot. conclusions play analysis of low-permeability reservoirs demands that the risk element of charge has to be treated in an unconventional manner. the low migration velocity of hydrocarbons in a low-permeability reservoir limits the areas of hydrocarbon reach. therefore, it is paramount to include a comprehensive analysis of the following factors to accurately define the play limits and increase the chances of making a discovery: 1. map of likely entry-points 2. analysis of pressure gradients 3. migration velocity 4. permeability evolution and variation in time and space. this innovative approach for semi-conventional plays would support a sound estimate of remaining resources in mature basins, and would constitute the best practice method in frontier exploration by diminishing the risk of failure. acknowledgements the author would like to express his gratitude to finn engstrøm and other colleagues at maersk oil (now total) and geus for very fruitful and challenging discussions when exploring for hydrocarbons in chalk and during the preparation of this article. john hopper from geus is 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chalk play: examples from the danish central graben. geological society, london, special publications 67, 247–282. https://doi.org/10.1144/gsl.sp.1992.067.01.10 megson, j.b. & tygesen, t. 2005: the north sea chalk: an underexplored and underdeveloped play. geological society, london, petroleum geology conference series 6, 159–168. https://doi.org/ 10.1144/0060159 milkov, a.v. 2015: risk tables for less biased and more consistent estimation of probability of geological success (pos) for segments with conventional oil and gas prospective resources. earth-science reviews 150, 453–476. https://doi.org/10.1016/j.earscirev.2015.08.006 https://doi.org/10.34194/geusb.v44.4836 http://www.geusbulletin.org https://doi.org/10.2118/71322-ms https://doi.org/10.2118/71322-ms https://doi.org/10.2118/114168-ms https://doi.org/10.2118/114168-ms https://doi.org/10.1016/j.energy.2004.03.071 https://doi.org/10.1016/s0928-8937(00)80016-8 https://doi.org/10.1144/0060145 https://doi.org/10.1144/0060145 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https://doi.org/10.1144/pgc8.24 vejbæk, o.v. et al. 2005: the history of hydrocarbon filling of danish chalk fields. geological society, london, petroleum geology conference series 6, 1331–1345. https://doi.org/10.1144/0061331 white, d.e. 1988: oil and gas play maps in exploration and assessment: geologic note 1. aapg bulletin 72, 944–949. https://doi. org/10.1306/703c911d-1707-11d7-8645000102c1865d winefield, p., gilham, r. & elsinger, r. 2005: plumbing the depths of the central graben: towards an integrated pressure, fluid and charge model for the central north sea hpht play. geological society, london, petroleum geology conference series 6, 1301–1315. https://doi. org/10.1144/0061301 https://doi.org/10.34194/geusb.v44.4836 http://www.geusbulletin.org https://doi.org/10.2118/31062-pa� https://doi.org/10.3997/1365-2397.2008013� https://doi.org/10.3997/1365-2397.2008013� https://doi.org/10.1144/0016-76492011-136� https://doi.org/10.1144/0016-76492011-136� https://doi.org/10.1144/pgc8.24� https://doi.org/10.1144/0061331� https://doi.org/10.1306/703c911d-1707-11d7-8645000102c1865d� https://doi.org/10.1306/703c911d-1707-11d7-8645000102c1865d� https://doi.org/10.1144/0061301� https://doi.org/10.1144/0061301� geological survey of denmark and greenland bulletin 6, 41-56 41geological survey of denmark and greenland bulletin 6, 41–56 © geus, 2004 the caledonian thin-skinned thrust belt of kronprins christian land, eastern north greenland a.k. higgins, n.j. soper, m. paul smith and jan a. rasmussen kronprins christian land in the extreme north of the east greenland caledonides, exposes a thin-skinned thrust belt up to 50 km wide developed in ordovician–silurian platform limestones and dolostones of the iapetus passive margin. this thrust belt is characterised by a series of ssw–nne-trending and east-dipping caledonian thrusts with westward displacements of generally a few kilometres each. it passes westwards into undisturbed autochthonous foreland. based on a line and area restoration, total displacement along a well-exposed wnw–ese section through the thrust belt amounts to 17.6 km, which represents a shortening of 45% in the line of section. biostratigraphic control in the limestone and dolostone succession is based on conodonts and macrofossils. the alteration colours of the conodonts provide estimates of maximum burial temperatures, which show that the thickness of the overlying thrust sheets ranged from about 6 to 12.5 km from west to east across the thrust belt. since the estimated former thickness of the vandredalen thrust sheet above the thin-skinned parautochthonous thrust belt is insufficient to yield the temperatures attained, higher thrust sheets must once have extended across the region. keywords: caledonides, conodonts, greenland, ordovician, thrust tectonics a.k.h. & j.a.r., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: akh@geus.dk n.j.s., gams bank, threshfield, skipton bd23 5np, uk. also affiliated with: department of geology, university college, galway, ireland. m.p.s., lapworth museum, school of geography, earth and environmental sciences, university of birmingham, edgbaston, birmingham b15 2tt, uk. the east greenland caledonides extend for 1300 km along the coastal region of east greenland between latitudes 70° and 82°n, in a belt up to 300 km wide. it can be broadly divided into an eastern thick-skinned thrust belt, and a western marginal thrust belt that in places is thin-skinned (fig. 1). the western marginal thrust belt is characterised by the presence of foreland windows, in most of which a thin lower palaeozoic sequence is preserved beneath the bordering thrusts demonstrating that the thrusting episode is postordovician (higgins et al. 2001a). the thrust sheets overlying the foreland windows incorporate substantial units of reworked basement gneisses, derived from the thick-skinned thrust belt to the east. the greenland inland ice obscures the western parts of the marginal thrust belt along most of its length, and the transition between the caledonian orogenic belt and the autochthonous foreland is only completely exposed in kronprins christian land (79°30′–82°n). here the transition zone takes the form of a thin-skinned parautochthonous thrust belt, which is the subject of this paper. geus bulletin 6.pmd 10-02-2005, 09:5441 42 ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ c a le d o n id e s greenland 70ºn 74ºn 78ºn 82ºw 35ºw wandel sea centrumsø 100 km jameson land scoresby sund 25 w traill ø wollaston forland bessel fjord danmarkshavn lambert land peary land station nord kr on pr ins c hr ist ian l an d nørreland window kronprins christian land thin-skinned thrust belt (parautochthonous foreland) palaeogene basalts palaeogene intrusions wandel sea basin: carboniferous–palaeogene sediments east greenland basins: carboniferous–cretaceous sediments devonian – continental sediments late to post-kinematic granites neoproterozoic–ordovician sediments (east greenland) neoproterozoic–silurian sediments (eastern north greenland) palaeo-mesoproterozoic sediments and basalts (eastern north greenland) crystalline complexes and sediments (archaean–mesoproterozoic) neoproterozoic–silurian sediments (north greenland) palaeo-mesoproterozoic sediments and basalts (eastern north greenland) mainly crystalline rocks – parautochthonous windows thrust fault/shear zone tectonic zone boundary post-caledonian late to post-caledonian caledonian orogenic belt caledonian foreland ▲ ▲ fig. 2 hamberg gletscher foreland dronning louise land charcot land window målebjerg window gåseland window c a le d o n ia n so le t h ru st in la n d ic e m a r g in a l t h r u s t b e lt t h ic k s k in n e d t h r u s t b e lt eleonore sø window ▲ ▲ ▲ fig. 1. general geological map of the east greenland caledonides, illustrating the division into western marginal and eastern thick-skinned thrust belts (modified from higgins et al. 2001a). the main foreland windows along the length of the fold belt are shown, with the kronprins christian land area in the extreme north. the frame indicates the area of fig. 2. geus bulletin 6.pmd 10-02-2005, 09:5442 43 geological setting throughout most of its length, the east greenland caledonides are dominated by crystalline orthogneiss complexes (fig. 1) that retain much of their ‘basement’ character despite caledonian reworking. the protolith age of the orthogneisses has been determined as archaean or proterozoic on the basis of numerous isotopic ages (e.g. steiger et al. 1979; kalsbeek et al. 1993, 1999). isotopic mineral ages are generally caledonian, testifying to widespread mediumto highgrade caledonian metamorphism (e.g. dallmeyer & strachan 1994; dallmeyer et al. 1994; brueckner et al. 1998). proterozoic sedimentary successions overlying the crystalline gneiss complexes are widespread in the southern half of the caledonides. an older late mesoproterozoic to early neoproterozoic succession (krummedal supracrustal sequence and equivalents; higgins 1988) preserves isotopic evidence of a pre-caledonian (~ 930 ma) thermal event, and in many areas hosts ~ 930 ma augen granite intrusions (jepsen & kalsbeek 1998; kalsbeek et al. 2000; watt et al. 2000; leslie & nutman 2000, 2003). the younger, neoproterozoic, eleonore bay supergroup is conspicuous in the fjord region of east greenland (72°–74°30′n), where it is unconformably overlain by the vendian tillite group and lower palaeozoic sediments, forming a succession up to 18.5 km thick. both sedimentary successions are variably affected by caledonian metamorphism and deformation, and both host caledonian granites (kalsbeek et al. 2001a, b). in the northern part of the east greenland caledonides, latest palaeoproterozoic to mesoproterozoic supracrustal successions are represented by the independence fjord group and associated volcanic rocks (figs 1, 2; see also below). these are widely exposed in the caledonian foreland west of danmark fjord, and are also conspicuously developed within the caledonian thrust complexes of kronprins christian land, where they are overlain by the neoproterozoic rivieradal group siliciclastic succession and hagen fjord group (fig. 2; see also stratigraphy section below). early work in southern kronprins christian land by fränkl (1954, 1955) established many of the principal structural features of this part of the east greenland caledonides. while subsequent interpretations of the frontal thrust systems were explained by hurst & mckerrow (1981a, b, 1985) in terms of three nappes, later systematic survey work has considerably simplified this view. the vandredalen thrust sheet is now recognised as the westernmost major allochthonous tectonicunit along the entire > 200 km long thrust front in kronprins christian land (fig. 2; rasmussen & smith 1996). the vandredalen thrust displaces the neoproterozoic rift succession now known as the rivieradal group(smithetal.2004a, this volume) across the parautochthonous foreland succession (higgins et al. 2001b). the thin-skinned thrust belt west of, and structurally underlying, the vandredalen thrust sheet is developed in an ordovician to lower silurian succession, that continues westwards into the undisturbed foreland sequences west of danmark fjord. the succession in this 30–50 km wide, parautochthonous thrust belt is disrupted by a series of east-dipping and ssw–nne-trending thrusts and associated belts of folding (fig. 2). a thin-skinned deformation style was also suggested in the earliest studies by fränkl (1954, 1955), and peel (1980) distinguished numerous significant thrusts in an w–e traverse through the belt in kronprins christian land west of romer sø. observations by peel indicated that the westernmost thrusts extend almost to danmark fjord. regional mapping of the southern part of kronprins christian land, including the parautochthonous thrust belt, was carried out during the 1993–1995 expeditions by the former geological survey of greenland (ggu; henriksen 1994a, b, 1995, 1996; higgins 1995). the vandredalen thrust climbs a steep ramp along the hekla sund – spærregletscher lineament, that is well exposed at the bay marmorvigen (m on fig. 2), is almost continuously exposed along the west side of hekla sund and extends northwards to the east side of brede spærregletscher (bs on fig. 2). west of the ramp, the thrust follows a long flat in the ordovician wandel valley formation, that is continuously exposed along the west side of sæfaxi elv, the river draining into marmorvigen. the > 200 km long vandredalen thrust sheet front has a general ssw–nne trend, and is traceable from west of blåsø through the east end of centrumsø to west of romer sø (fig. 2). this trend line coincides with another ramp that cuts up through the ordovician–silurian platform limestones and dolomites and carries the vandredalen thrust sheet up to overlie silurian turbidites of the lauge koch land formation at present-day exposure levels. the root zone of the vandredalen thrust sheet, along the ssw–nne-trending hekla sund – spærregletscher lineament, coincides approximately with the west margin of the original rift basin (hekla sund basin) in which the rivieradal group succession accumulated (higgins et al. 2001b). geus bulletin 6.pmd 10-02-2005, 09:5443 44 wandel sea basin sequence (post-caledonian) samuelsen høj formation lauge koch land formation odins fjord formation turesø formation wandel valley formation kap holbæk formation crystalline basement thrust fault børglum river and sjælland fjelde formations fyns sø, kap bernhard, campanuladal fms hagen fjord gp rivieradal group independence fjord gp and basaltic formations vandredalen thrust hagen fjord group   vandredalen thrust sheet   ▲ ▲ ▲ ▲ ro mer sø d an m ar k f j o r d amdrup land holm land hovgaard ø kap bernhoft dijmphna sund centrumsø sk a l l in g e n syd vej dal rivieradal 20°w 81°n blåsø nioghalvfje rdsfjorden sæ faxi elv h ek la su nd ingolf fjo rdbs h fl m va nd re da le n ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ 80°n sp t sp t ▲ ▲ c a le d o n id es greenland 25 km geus bulletin 6.pmd 10-02-2005, 09:5444 45 east of the hekla sund – spærregletscher lineament a broad zone of latest palaeoproterozoic to mesoproterozoic clastic and volcanic rocks crops out, and still farther east crystalline basement rocks extend to the eastern coast of kronprins christian land (fig. 2). these broad regions are bounded by steeply inclined shear zones, some of which probably represent major thrusts. the crystalline basement rocks underlying the post-caledonian wandel sea basin succession in the coastal zone incorporate eclogitic enclaves that testify to deep burial during the caledonian orogeny, followed by rapid exhumation (e.g. gilotti & ravna 2002; gilotti et al. 2003). the pronounced ssw–nne lineament that can be traced from hovgaard ø through western holm land to amdrup land, is generally not well exposed, but appears to have a complex history. this feature is often viewed as a northward continuation, or a splay, of the major, sinistral, storstrømmen shear zone, described from hertugen af orléans land (78°n) by strachan & tribe (1994). the latest movements on the lineament in kronprins christian land are post-caledonian, with eastward downthrow of the wandel sea basin succession. however, in southern hovgaard ø and lambert land jones & escher (1995) record a series of late caledonian ductile shear zones along the lineament, that preserve evidence of both sinistral and east-side-up displacement. in lambert land these shear zones post-date foreland-propagating thrust-stacking events, that place thrust sheets of high-grade crystalline gneisses (with eclogitic enclaves) above thrust sheets comprising independence fjord group sandstones. it is considered likely that the high grade basement gneisses of hovgaard ø and holm land form part of major, thick-skinned thrust sheets that once projected westwards, structurally above the strongly sheared and folded independence fjord group west of the hovgaard ø – amdrup land lineament (see also fig. 5). stratigraphy the autochthonous and parautochthonous foreland comprises thick latest palaeoproterozoic to mesoproterozoic successions (hekla sundformation, aage berthelsen gletscher formation, independence fjord group, zig-zag dal basalt formation) and associated mafic intrusions (midsommersø dolerite formation); see also sønderholm & jepsen (1991). these are overlain by neoproterozoic shelf sediments (hagen fjord group: comprising the jyske ås, campanuladal, kap bernhard and fyns sø formations). there is a hiatus between the fyns sø formation dolostones and the overlying sandstones of the kap holbæk formation with local developments of palaeokarst (smith et al. 1999). another hiatus occurs between the kap holbæk formation (early cambrian) and the overlying ordovician–silurian carbonate and siliciclastic rocks. the neoproterozoic rivieradal group is represented only in the allochthonous vandredalen thrust sheet, and its deposition can be linked to an episode of extensional rifting (higgins et al. 2001b). all these units were involved to some extent in the caledonian folding and thrusting, but in the thin-skinned parautochthonous belt the thrusts are essentially confined to the ordovician–silurian sequence. the best exposed sections through the thin-skinned thrust belt follow the sides of centrumsø and the valleys which branch off the west end of this lake. this is the only area where there is sufficient relief and ground control to permit reconstruction of a restorable section (see below). other good partial sections occur in valleys to the north and south. the extensive plateau areas between valleys are often poorly exposed, and here mapping was carried out by spot checks of the sporadic exposures, supplemented by sampling and conodont studies (rasmussen & smith 2002). the main formations represented on the maps and cross-sections are listed in fig. 3, and are briefly described below. hekla sund formation, aage berthelsen gletscher formation, independence fjord group, midsommersø dolerite formation, zig-zag dal basalt formation with the exception of the tholeiitic basalts of the hekla sund formation and aage berthelsen gletscher formation, the type areas of these proterozoic divisions were established on the caledonian foreland west of facing page: fig. 2. geological map of southern and central kronprins christian land. the frame outline centred on centrumsø indicates the position of the cross-section and geological map presented in fig. 4; the extension of the section line beyond the frame is that of the cross-section in fig. 5. the lineament traceable from hovgaard ø through western holm land to amdrup land, marked by a dashed line, has a complex history (see text). bs, brede spærregletscher; fl, finderup land; h, hjørnegletscher; m, marmorvigen; spt, spærregletscher thrust. modified from higgins et al. (2001a). geus bulletin 6.pmd 10-02-2005, 09:5445 46 danmark fjord (sønderholm & jepsen 1991). the independence fjord group comprises a more than 2 km thick succession of mainly clastic alluvial deposits, dominantly white-weathering quartzitic sandstones (collinson 1980, 1983). the midsommersø dolerite formation consists of the widespread doleritic sheets, sills and dykes which invade the independence fjord group sandstones (jepsen 1971; kalsbeek & jepsen 1983). the zig-zag dal basalt formation comprises at least 1350 m of lava flows which overlie the independence fjord group (jepsen et al. 1980; kalsbeek & jepsen 1984); they are considered to be the extrusive equivalents of the midsommersø dolerite formation. highly deformed quartzitic sandstones, doleritic dykes and basaltic lava sequences which crop out in the alpine region of eastern kronprins christian land within the caledonian orogenic belt have traditionally been regarded as equivalents of the foreland divisions. however, the survey’s regional mapping revealed that the basaltic sequences found in the thrust complexes of kronprins christian land do not overlie the independence fjord group, but are interbedded with the lower levels of the quartzite succession. these basalfig. 3. summary stratigraphic scheme of proterozoic and palaeozoic units depicted on the maps, and their relationships to iapetus opening (modified from smith et al. 1999). non-deposition or erosion is depicted by vertical ruling. silurian ordovician cambrian vendian sturtian riphean lauge koch land formation samuelsen høj formation odins fjord formation turesø formation børglum river formation sjælland fjelde formation wandel valley formation kap holbæk formation hagen fjord group fyns sø fm kap bernhard fm campanuladal fm jyske ås fm zig-zag dal basalt formation independence fjord group hekla sund fm, aage berthelsen gletscher fm, & interbedded quartzites thermal subsidence extensional rifting and block tilting extensional riftingrivieradal group (allochthonous vandredalen thrust sheet only) post-rift thermal subsidence baltica collision thrust loaded flysch basin tectonic setting depositional environment stratigraphy lapetus passive margin lapetus opening pre-lapetus rift-sag cycle intracratonic extensional events thermal subsidence block tilting zz if hs/ab kh rg    geus bulletin 6.pmd 10-02-2005, 09:5446 47 tic sequences are distinguished as the hekla sund formation and aage berthelsen gletscher formation (pedersen et al. 2002). shrimp isotopic studies on rhyolites of the hekla sund formation yielded an age of 1740 ma (kalsbeek et al. 1999). this result implies that either the age range of the independence fjord group must be extended downwards to the later part of the palaeoproterozoic, or there are two superficially indistinguishable quartzite sequences, of which the older unnamed succession is interbedded with the hekla sund and aage berthelsen gletscher formations. the first alternative is adopted here. quartzitedyke-basalt associations similar to the foreland succession are presumed to underlie the entire parautochthonous region. rivieradal group the succession of sandstones, mudstones, conglomerates and some carbonate rocks first mapped by fränkl (1954, 1955), and assigned by hurst & mckerrow (1981a, b) to a single sequence that they referred to as the ‘rivieradal sandstones’, has been formally defined as the rivieradal group (smith et al. 2004a, this volume). the rivieradal group is restricted to the vandredalen thrust sheet, where it is overlain conformably by units of the hagen fjord group. fränkl had recognised that the neoproterozoic rivieradal group was not represented on the foreland, and introduced the term ‘hekla sund basin’ for its area of deposition. field work by ggu in 1993–1995 demonstrated that the rivieradal group is 7.5–10 km thick. it was also shown that the sediments of the rivieradal group had accumulated in an east-facing, half-graben rift-basin, bounded to the west by extensional faults; this basin was estimated to have been at least 200 km long and 50 km wide (higgins et al. 2001b). during the caledonian orogeny the rivieradal group was displaced westwards across the western margin of the rift basin as the vandredalen thrust sheet. the root zone of this thrust sheet and the remnants of the original rift basin can be traced in a narrow belt through the centre of the alpine region along the hekla sund – spærregletscher lineament (fig. 2). hagen fjord group (and kap holbæk formation) representatives of the hagen fjord group are preserved in the frontal portions of the vandredalen thrust sheet, resting conformably on the rivieradal group. in the foot wall of the vandredalen thrust, as elsewhere in eastern north greenland, the hagen fjord group rests directly on independence fjord group lithologies, locally with an intervening basal clastic unit. the hagen fjord group is thus viewed as a transgressive, post-rift sequence, and its presence in both the hanging wall and foot wall of the vandredalen thrust enables the displacement on the vandredalen thrust to be estimated at 35–50 km (higgins et al. 2001b). the jyske ås formation (fig. 3) at the base of the group occurs only west of danmark fjord in the foreland, and is not considered further here. in the parautochthonous region with which this paper is concerned, four formations are recognised in addition to the basal clastic unit, although the uppermost unit (the kap holbæk formation) is now formally excluded from the hagen fjord group (see below). 1. basal clastic unit. this dominantly conglomeratic unit directly overlies independence fjord group quartzitic sandstones, and is overlain by siltstones and mudstones ascribed to the campanuladal formation. the unit was first recorded at hjørnegletscher (h on fig. 2) on the north side of inner ingolf fjord (jepsen & kalsbeek 1981) where it is a few metres thick. in 1993 two additional developments of the unit, respectively 35 m and 0–60 m thick, were located along the margin of the alpine region north of sæfaxi elv (jepsen et al. 1994). 2. campanuladal formation. dominated by green and red fine-grained sandstones, siltstones and mudstones, it is about 110–175 m thick in the foreland areas west of the head of danmark fjord (clemmensen & jepsen 1992). in the parautochthonous region between inner ingolf fjord and sæfaxi elv, jepsen & kalsbeek (1985) reported 0–80 m of mudstone and sandstone of the formation overlying either the basal conglomeratic unit or the independence fjord group. 3. kap bernhard formation. this comprises reddishbrown limestones with minor amounts of silt, and is about 150 m thick at the head of danmark fjord (clemmensen & jepsen 1992). the formation is up geus bulletin 6.pmd 10-02-2005, 09:5447 48 to 400 m thick in the frontal region of the vandredalen thrust sheet. 4. fyns sø formation. at its type locality at the head of danmark fjord (craig & jepsen 1995), it is made up of 356 m of spectacular, cliff-forming, yellowweathering dolostones, characteristically preserving well-formed stromatolites. a similar thickness (~ 400 m) is seen in both the foot wall and the hanging wall of the vandredalen thrust. 5. kap holbæk formation. this was originally the upper formation of the hagen fjord group (clemmensen & jepsen 1992). recognition that the formation is early cambrian, and that the hiatus between it and the underlying fyns sø formation covers the entirevendian (fig. 3), led smith et al. (2004b, this volume) to formally exclude it from the hagen fjord group. the formation was recognised in the parautochthonous belt in the inner parts of ingolf fjord in 1994 (jepsen & sønderholm 1994), and here is up to 180 m thick; it comprises variegated mudstones at the base overlain by a light and dark coloured sandstone succession. sandstone-filled fissures and cave-like lenses in the upper surface of the underlying fyns sø formation, first recorded by fränkl (1954, 1955), have been interpreted as palaeokarst (smith et al. 1999). the kap holbæk formation was recognised in the hanging wall of the vandredalen thrust by hurst & mckerrow (1981a, b), who placed it in their ‘finderup land nappe’. lower palaeozoic platform the lower palaeozoic platform strata of eastern north greenland are the easternmost representatives of the franklinian basin succession, which is exposed in a broad, 900 km long belt across north greenland (higgins et al. 1991). the earliest lower palaeozoic platform strata in the parautochthonous belt of eastern kronprins christian land are the early ordovician limestones and dolostones of the wandel valley formation (rasmussen & smith 1996; smith et al. 2004b, this volume), which rest unconformably on the fyns søformation or kap holbæk formation. uplift of eastern north greenland and subsequent erosion have resulted in a progressive overstep of the early ordovician from west to east across north greenland (peel & smith 1988). there is also a north–south component to the overstep, since south of kronprins christian land, in lambert land, the hagen fjord group is missing and the wandel valley formation rests directly on independence fjord group lithologies (smith et al. 1999; smith 2000). a fuller stratigraphical description of the lower palaeozoic platform limestone anddolostonesuccession is givenbysmith et al. (2004b, this volume). the following formations are distinguished on the maps and cross-sections of this paper. 1. wandel valley formation (upper ibexian – middle whiterockian). three limestone and dolostone members are present in the parautochthonous belt, all very similar in their development to their counterparts on the foreland around danmark fjord, and with a total thickness of about 335 m. 2. sjælland fjelde formation (upper whiterockian). about 100 m thick, it is divided into a lower dark grey burrow-mottled limestone and dolostone unit and an upper grey dolostone unit. the vandredalen thrust follows a long flat in the middle part of the formation, well seen along the west side of sæfaxi elv, before climbing a ramp to another flat in the upper dolomite unit. near the head of ingolf fjord, about 70 km to the north, the vandredalen thrust occupies the same stratigraphic level. 3. børglum river formation (mohawkian – upper cincinatian). the formation is widespread in the parautochthonous belt, where it comprises a thick succession of dominantly dark, nodular, burrowmottled limestones with abundant fossils. a complete section through the unit is not seen in the parautochthonous belt, but is probably close to the thickness of 430 m measured in the autochthonous foreland areas further to the north-west (smith et al. 1989). facing page: fig. 4. geological map and restored cross-section of the thinskinned thrust belt in the centrumsø region. a: geological map and location of section line; see also frame in fig. 2. base camp indicated by filled triangle. b: cross-section with calculated displacements on individual thrusts in kilometres (e.g. 2.75) based on a line and area balance; only the sjælland fjelde formation is given a distinctive ornament, with other formations indicated by two-letter abbreviations (see legend on map of fig. 4a). note the gently eastwardsdipping floor thrust at the base of the wandel valley formation. c: model section with thrusts restored; note reproduced at a smaller scale than the cross-section in b. in both b and c the thrusts are indicated by thicker lines. geus bulletin 6.pmd 10-02-2005, 09:5448 49 ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ cen tru m sø section line 0 5 km 0 5 km 0 5 km n lauge koch land odins fjord turesø sjælland fjelde fyns sø kap bernhard rivieradal group thrust lk of tu br sf fs wv wv kb rg wandel valley børglum river ▲ ▲ base camp a b c lk of of tu br wv lk tu tu br br br br br wv wv br wv lk of tu lk lk of lk lk of oftu tubr br fs kb rg wv br of tu tu wv 0.05 0.55 0.4 0.9 0.95 0.751.8 1.0 2.85 5.4 km 1.0 km 2.75 2.75 4.9 0.8 9.15 km 8.45 km 22.4 km g e u s b ulletin 6.pm d 10-02-2005, 09:54 49 50 4. turesø formation (upper cincinnatian – lower llandovery). the formation spans the ordovician– silurian boundary (armstrong 1990), and where measured 7 km west of centrumsø comprises about 200 m of variably coloured dolostones and limestones (see fig. 6a). the colour variations make the formation conspicuous and easily recognisable. towards the eastern end of centrumsø, just west of the vandredalen thrust sheet front, the formation thickens to at least 350 m; here trains of tight folds are developed in the dolostone-dominated intervals (see fig. 6b). 5. odins fjord formation (mid-llandovery). the formation is widely exposed in southern kronprins christian land close to the vandredalen thrust sheet front, where it is at least 220 m thick, although deformation and poor exposure make this estimate uncertain. christie & peel (1977) estimated a thickness of 320 m in south-east peary land. the transition from the underlying turesø formation is marked by a change in colour from pale grey to pale brown, and in lithology from dolostone to limestone rich in tabulate corals and stromatoporoids. 6. samuelsen høj formation (upper llandovery). developed as conspicuous reefs, the formation is represented by several major bodies in northern kronprins christian land; those in southern areas are generally smaller and mainly occur in a belt just west of the vandredalen thrust sheet front (fig. 2). only one small body is known south of centrumsø, and there are none in the line of section (fig. 4). 7. lauge koch land formation (uppermost llandovery – wenlock). the silurian flysch of kronprins christian land was assigned by hurst & surlyk (1982) to fränkl’s ‘profilfjeldet shales’, which was given member rank within the lauge koch land formation. the sequence is widely involved in the major thrusts of northern kronprins christian land, where a maximum thickness of 400 m was estimated (hurst & surlyk 1982). further south the formation crops out mainly in a zone just west of the vandredalen thrust sheet front. only two thrust-bounded inliers occur south of centrumsø; here the lower 50 m of the formation is characterised by black shaly siltstones interbedded with dark grey to black bituminous and nodular carbonate rocks (smith et al. 2004b, this volume). structure the most important thrusts within the 30–50 km wide thin-skinned thrust belt of kronprins christian land are depicted in figs 2 and 4. they make up a major imbricate stack beneath a former extension of the vandredalen thrust sheet. individual thrusts dip eastwards at angles varying from about 30° to 70°, although the steeper thrusts appear to represent over-steepening arising from further thrust displacement in the foot wall succession. many thrusts can be followed for several tens of kilometres, some for as much as 75 km along strike. major folding accompanied the thrusting, although this is normally conspicuous only in certain formations. about 25 km south of centrumsø several major thrusts die out. further south, only one major thrust has been traced for 30 km west of the vandredalen thrust front, and this divides into two thrusts west of blåsø (fig. 2). all the ordovician–silurian stratigraphic units from the wandel valley formation to the lauge koch land formation are involved in the thrusting. individual thrust movements range from a few hundred metres to several kilometres. despite topographic relief of 1000 m, matching foot wall and hanging wall cut-offs are rarely observed. estimates of thrust displacements therefore rely on the construction of a restorable crosssection. the best exposed sections are, in the north, the valley system west of romer sø (described by peel 1980) and, in the south, the valley system containing centrumsø with which this paper is mainly concerned. a restorable cross-section through the centrumsø area constructed perpendicular to the thrust trends is presented in fig. 4. prior to attempting to restore the cross-section, a series of cross-sections (not reproduced here) were constructed along profile lines north and south of centrumsø to gain an impression of the possible range of displacements. the section line of fig. 2 was chosen because of the excellent exposures on the cliff walls north and south of the lake, and because of the generally good ground control. initial section construction was at a scale of 1:50 000, on the basis of enlarged copies of the survey’s 1:100 000 topographic maps. thrust trajectories and fold shapes were projected into the line of section using the best available thickness estimates for formations as noted above. the maximumobserved thicknessestimateof400 m for the lauge koch land formation was used. in respect of the turesø formation, the 200 m thickness was used in the west, and 350 m in the eastern part of the section. geus bulletin 6.pmd 10-02-2005, 09:5450 51 balancing was attempted initially assuming that a single floor thrust in the parautochthonous belt followed the base of the børglum river formation (as the wandel valley and sjælland fjelde formations were not visibly involved in the thrusting along the line of profile). however, all attempts at a balance with this constraint produced an unrealistic undulating floor thrust (not illustrated here). the floor thrust was then reassigned downwards to the base of the wandel valley formation with, in addition, a major thrust at the base of the børglum river formation. this change is justified on the grounds that: (1) the wandel valley and sjælland fjelde formations are both involved in the thrusting in northern kronprins christian land (see fig. 1 in peel 1980); (2) along sæfaxi elv, immediately north of the eastward extension of the centrumsø cross-section, several highly disturbed bedding-parallel shear zones were observed near the base of the wandel valley formation. the restoration of the centrumsø cross-section achieved on this basis (fig. 4b), exhibits a very gentle eastward inclination for the floor thrust at the base of the wandel valley formation. in this model all the thrusts west of the end of centrumsø root into the floor thrust, whereas the thrusts exposed along the margins of centrumsø all root into the slightly higher flat thrust following the base of the børglum river formation. the section restoration presented in fig. 4 involved resolution of several problems. at the west end of the section there is a very broad mapped expanse of børglum river formation between two observed major thrusts (see map fig. 4a). the 4.5 km long valley section shows the sequence dipping at moderate angles eastwards, with locally some dislocation and associated folding. however, since the maximum thickness of the børglum river formation is probably about 430 m, restoration could only be achieved assuming the formation to be repeated in a duplex with displacements of 400–1800 m on the individual thrusts. the positions of the duplex thrusts were not identified during the field work, mainly because the significance of the over-thickened section was not appreciated; thus, while depicted on the cross-sections (fig. 4b, c), these thrusts are not shown on the map (fig. 4a). a further problem concerned a long central segment of the cross-section which exhibits a syncline at the west end (see fig. 6a) and a broad flat anticline in the centre with the lowest levels of the børglum river formation exposed at valley level. this could only be satisfactorily accommodated by introducing a ramp duplicating the wandel valley and sjælland fjelde formations over a distance of 5.4 km (central part of section in fig. 4b). in the cliff north of the base camp at the west end of centrumsø, a long flat thrust brings the børglum river formation above a thin sequence of the turesø formation. south of centrumsø the same thrust changes levels and takes the turesø formation above the odins fjord formation on an equally long flat thrust. similar long thrust flats are interpreted to exist at the eastern end of the section, with the largest displacement on an individual thrust estimated at 4.9 km. intense folding at the eastern end of the section, just west of the vandredalen thrust sheet front (see fig. 6b), and the implications of such internal distortion in other parts of the section, cannot be accurately depicted. total displacement on the basis of the model restoration in fig. 4b is estimated at 17.6 km. the thrusts depicted in the east part of the section along centrumsø have a total displacement of 8.45 km rooting into the thrust at the base of the børglum river formation, which merges with the vandredalen thrust at the vandredalen thrust front. the thrusts west of centrumsø root into a floor thrust following the base of the wandel valley formation and have an estimated total of 9.15 km displacement; this thrust merges with the vandredalen thrust east of the line of section in the vicinity of marmorvigen (fig. 2). the restoration implies that an original 43 km wide segment of the parautochthonous belt has been reduced to about 25.4 km in the line of section, a shortening of approximately 45%. the restored section depicted in fig. 4b demonstrates that the model chosen is realistic. it invokes only two major flat thrusts, both of which merge eastwards with the vandredalen thrust. conodont geothermometry epstein et al. (1977) demonstrated that colour variations of conodont elements are principally related to temperature. they erected a scale of conodont alteration indices (cai 1–5) ranging from pale yellow through shades of brown to black, corresponding to a temperature range from < 50°–300°c. higher alteration indices (cai 6–8), in which the conodont elements progressed from black through grey to white, were calibrated by rejebian et al. (1987) as corresponding to a temperature range from 300°c to over 600°c. a regional description of conodont geothermometry geus bulletin 6.pmd 10-02-2005, 09:5451 52 in the kronprins christian land area has been presented by rasmussen & smith (2001). conodonts studied in kronprins christian land were recovered from stratigraphic units ranging in age from mid-early ordovician (wandel valley formation) to llandovery (lauge koch land formation). lithologies varied from unaltered platform dolostones and limestones to their highly sheared equivalents underlying the vandredalen thrust sheet. whereas the degree of internal shearing and deformation had a significant effect on the morphological character of the conodont elements, it had no apparent effect on the colour alteration indices. the cai isothermal zones run parallel to the thrust trends and the vandredalen thrust sheet front in southern kronprins christian land. cai values of 2–3 were seen west of danmark fjord. a broad zone of cai 3 extends eastwards to approximately the west limit of the cross-section in fig. 4. most of the cross-section is within the zone of cai 4, rising to cai 5 at the eastern end adjacent to the front of the vandredalen thrust sheet. the limestones and dolostones beneath the vandredalen thrust sheet, exposed along sæfaxi elv, are in cai zone 5 increasing to cai 5–6 in the easternmost exposures at marmorvigen. the cai temperatures indicate the maximum thickness of the caledonian overburden, comprising the vandredalen thrust sheet and possible higher thrust units. the thickness was determined from estimates of geothermal gradients and the thermal conductivity of the rock units involved (see rasmussen & smith 2001, for details). the results imply that the approximate thickness of the maximum overburden in the area of the cross-section (fig. 5), ranged from about 6 km at the west end of the cross-section to 10.7 km farther east at the front of the vandredalen thrust sheet (fig. 6c). the highest cai values at marmorvigen point to an overburden of 12.5 km (rasmussen & smith 2001). the assumed extent and thickness of the vandredalen thrust sheet formerly present above the parautochthonous zone are also indicated in fig. 5. the hagen fjord group in the hanging wall exhibits a cut-off against the vandredalen thrust along much of the vandredalen thrust sheet front on the west side of vandredalen. thus, the former extent of the vandredalen thrust sheet across the parautochthonous zone must have consisted essentially of a packet of ordovician–silurian carbonate and siliciclastic rocks. the thickness of this packet was probably not much greater than 2 km (fig. 5). the only uncertainty in this estimation of the thickness concerns the contribution of turbidites of the lauge koch land formation. a maximum thickness of 400 m has been assumed for this unit in the cross-section, being the maximum thickness preserved in present-day exposures (hurst & surlyk 1982). the silurian turbidites of north greenland were derived from erosion of the rising caledonian mountain chain. the thickness of the turbidite succession that may have accumulated in the western part of present kronprins christian land before it was over-ridden by the westward-propagating caledonian thrust sheets is unknown. between 3 and 10 km of additional overburden above the vandredalen thrust 0 km 5 0 22 km 6 km 12.5 km 50 km ? 100 km ? ord.–sil st 18 km wsw ese vt spt 50 km es timated ov erburden f ro m con odont a lt erat i o n ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ crystalline basement rivieradal group hagen fjord group ordovician – silurian independence fjord group and associated volcanics fig. 5. simplified cross-section through the caledonian fold belt in kronprins christian land, from higgins et al. (2001b); for section line see fig. 2. the maximum overburden deduced from conodont alteration indices (indicative of eastward increase in temperature) is also shown. spt, spærregletscher thrust; st, caledonian sole thrust; vt, vandredalen thrust. geus bulletin 6.pmd 10-02-2005, 09:5452 53 of of tu tu rg vt a b c fig. 6. a: syncline in line of cross-section looking north, 7 km west of the centrumsø base camp. of, odins fjord formation; tu, turesø formation. cono-donts have cai values of 4, indicative of a former overburden of about 6.8 km. summit at centre is 500 m above the valley floor. photo: j. lautrup. b: intense folding in variegated dolomites of the turesø formation (tu). north side of centrumsø, about 3 km west of the vandredalen thrust front. conodonts have cai values of 4–5, indicative of an overburden of about 8–9 km. plateau is about 750 m above the lake level (foreground). photo: j. lautrup. c: outlier of rivieradal group (rg) conglomerates and sandstones in the vandredalen thrust sheet, overlying ordovician carbonates of the odins fjord formation (of) on the west side of vandredalen. the vandredalen thrust (vt) follows the marked discordance. conodonts from the carbonates of the odins fjord formation have cai values of 5, indicative of a former overburden of about 10.7 km. summit is 850 m above the valley floor in the foreground. geus bulletin 6.pmd 10-02-2005, 09:5453 54 sheet would be required to reach the temperatures demonstrated by the conodont alteration pattern, and it is considered unlikely that this can be accounted for by substantially increasing only the contribution of the lauge koch land formation turbidites. it is more probable that higher, westward-propagating thrust sheets were formerly present above the vandredalen thrust sheet. these are likely to have comprised units such as the independence fjord group quartzitic sandstones with associated dolerite dykes and sills, and the hekla sund formation basalts (representatives of which crop out in the mountainous region east of the hekla sund – spærregletscher lineament). these units would have been transported westwards on the spærregletscher thrust (spt in figs 2, 5). lower palaeozoic formations may also have been present in the proximal parts of this thrust sheet. all the cai zones are based on sample collections from the parautochthonous zone structurally underlying the vandredalen thrust sheet. this zone is part of a thin-skinned thrust belt, and therefore the most likely setting to account for the increased temperatures would be burial of the parautochthonous zone beneath a pile of westward-directed caledonian thrust sheets. allowing for subsidence of the parautochthonous zone that resulted from the weight of the overlying thrust burden, the thrust sheets must still have made up a substantial mountain chain, increasing in altitude eastwards where summits may have attained altitudes of about 3–4 km. conclusions the rock units which constitute the up to 50 km wide thin-skinned thrust belt west of the vandredalen thrust front extend westwards into undisturbed foreland. the thrust belt is therefore viewed as parautochthonous. the deformation associated with the eastward-dipping thrusts of the parautochthonous zone involves only ordovician–silurian rock units and is essentially thinskinned in style. a line-and-area restoration along the best exposed section through the thrust belt, i.e. along centrumsø and adjacent valleys, can be achieved assuming that the observed thrusts root into two flat thrusts. one is depicted as the caledonian floor or sole thrust, in this region located at the base of the wandel valley formation; the second slightly higher thrust is assumed to lie at the base of the børglum river formation. both thrusts are assumed to merge eastwards with the vandredalen thrust. total displacement of 17.6 km on the two flat thrusts in the model restoration implies that an original 43 km wide segment of the parautochthonous belt has been reduced to 25.4 km, a shortening of 45% in the line of section. the colour changes experienced by conodont elements reflect variations in temperature which can be linked to the maximum thickness of overburden during the caledonian orogeny. overburden estimates increase systematically from 6 km at the west end of the cross-section to 10.7 km at the vandredalen thrust front, and farther east to 12.5 km at marmorvigen (fig. 5). as the vandredalen thrust sheet overlying the parautochthonous zone was probably not much more than 2 km thick, the remainder of the estimated overburden must have comprised higher thrust sheets, since eroded, that projected westwards across the parautochthonous belt. acknowledgements the field and conodont studies of j.a.r. were financially supported by the carlsberg foundation (grant no. 950164/20-1292). the helpful comments of the two reviewers, robin strachan and brian chadwick, are gratefully acknowledged. references armstrong, h.a. 1990: conodonts from the upper ordovician – lower silurian carbonate platform of north greenland. bulletin grønlands geologiske undersøgelse 159, 151 pp. brueckner, h.k., gilotti, j.a. & nutman, a.p. 1998: caledonian eclogite-facies metamorphism of early proterozoic protoliths from the north-east greenland eclogite province. contributions to mineralogy and petrology 130, 103–120. christie, r.l. & peel, j.s. 1977: cambrian–silurian stratigraphy of børglum elv, peary land, eastern north greenland. rapport grønlands geologiske undersøgelse 82, 48 pp. clemmensen, l.b. & jepsen, h.f. 1992: lithostratigraphy and geological setting of upper proterozoic shoreline-shelf deposits, hagen fjord group, eastern north greenland. rapport grønlands geologiske undersøgelse 157, 27 pp. collinson, j.d. 1980: stratigraphy of the independence fjord group (proterozoic) of eastern north greenland. rapport grønlands geologiske undersøgelse 99, 7–23. collinson, j.d. 1983: sedimentology of unconformities within a fluvio-lacustrine sequence; middle proterozoic of eastern north greenland. sedimentary geology 34, 145–166. craig, l.e. & jepsen, h.f. 1995: late proterozoic hagen fjord group: stratigraphic relationships between the fyns sø forgeus bulletin 6.pmd 10-02-2005, 09:5454 55 mation and kap holbæk formation – including a new reference section through the fyns sø formation. in: higgins, a.k. (ed.): express report: eastern north greenland and northeast greenland 1995, 53–57. unpublished report, geological survey of greenland, copenhagen. dallmeyer, r.d. & strachan, r.a. 1994: 40ar/39ar mineral age constraints on the timing of deformation and metamorphism, north-east greenland caledonides. in: higgins, a.k. (ed.): geology of north-east greenland. rapport grønlands geologiske undersøgelse 162, 153–162. dallmeyer, r.d., strachan, r.a. & henriksen, n. 1994: 40ar/39ar mineral age record in ne greenland: implications for tectonic evolution of the north atlantic caledonides. journal of the geological society (london) 151, 615–628. epstein, a.g., epstein, j.b. & harris, l.d. 1977: conodont color alteration – an index to organic metamorphism. professional paper united states geological survey 995, 27 pp. fränkl, e. 1954: vorläufige mitteilung über die geologie von kronprins christians land (ne-grönland, zwischen 80°–81°n und 19°–23°w). meddelelser om grønland 116(2), 85 pp. fränkl, e. 1955: weitere beiträge zur geologie von kronprins christians land (ne-grönland, zwischen 80° und 80°30′n). meddelelser om grønland 103(7), 35 pp. gilotti, j.a. & ravna, e.j.k. 2002: first evidence for ultrahighpressure metamorphism in the north-east greenland caledonides. geology 30, 551–554. gilotti, j.a., nutman, a.p, brueckner, h.k. & mcclelland, w.c. 2003: devonian collision along the northeastern margin of laurentia, greenland caledonides. geological society of america: abstracts with programs 35, n. 3. henriksen, n. 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(eds): sedimentary basins of north greenland. bulletin grønlands geologiske undersøgelse 160, 71–139. higgins, a.k., leslie, a.g., smith, m.p. & rasmussen, j.a. 2001a: neoproterozoic – lower palaeozoic stratigraphical relationships in the marginal thin-skinned thrust belt of the east greenland caledonides: comparisons with the foreland in scotland. geological magazine 138(2), 143–160. higgins, a.k., soper, n.j., leslie, a.g., smith, m.p., rasmussen, j.a. & sønderholm, m. 2001b: the neoproterozoic hekla sund basin, eastern north greenland: a pre-iapetan extensional sequence thrust across its rift shoulders during the caledonian orogeny. journal of the geological society (london) 158, 487–489. hurst, j.m. & mckerrow, w.s. 1981a: the caledonian nappes of eastern north greenland. nature, london 290, 772–774. hurst, j.m. & mckerrow, w.s. 1981b: the caledonian nappes of kronprins christian land, eastern north greenland. rapport grønlands geologiske undersøgelse 106, 15–19. hurst, j.m. & mckerrow, w.s. 1985: origin of the caledonian nappes of north-east greenland. in: gee, d.g. & sturt, b.a. (eds): the caledonide orogen – scandinavia and related areas, 1065–1069. chichester: john wiley & sons. hurst, j.m. & surlyk, f. 1982: stratigraphy of the silurian turbidite sequence of north greenland. bulletin grønlands geologiske undersøgelse 145, 121 pp. jepsen, h.f. 1971: the precambrian, eocambrian and early palaeozoic stratigraphy of the jørgen brønlund fjord area, peary land, north greenland. meddelelser om grønland 192(2), 42 pp. jepsen, h.f. & kalsbeek, f. 1981: non-existence of the carolinidian orogeny in the prinsesse caroline-mathilde alper of kronprins christian land, eastern north greenland. rapport grønlands geologiske undersøgelse 106, 7–14. jepsen, h.f. & kalsbeek, f. 1985: evidence for non-existence of a carolinidian fold belt in eastern north greenland. in: gee, d.g. & sturt, b.a. (eds): the caledonide orogen – scandinavia and related areas, 1071–1076. chichester: john wiley & sons. jepsen, h.f. & kalsbeek, f. 1998: granites in the caledonian fold belt of east greenland. in: higgins, a.k. & frederiksen, k.s. (eds): caledonian geology of east greenland 72°–74°n: preliminary reports from the 1997 expedition. danmarks og grønlands geologiske undersøgelse rapport 1998/28, 73–82. jepsen, h.f. & sønderholm, m. 1994: sedimentological studies of the hagen fjord group and ‘rivieradal sandstones’ (late proterozoic), eastern north greenland. in: henriksen, n. (ed.): express report: eastern north greenland and north-east greenland 1994, 39–48. unpublished report, geological survey of greenland, copenhagen. jepsen, h.f., kalsbeek, f. & suthren, r.j. 1980: the zig-zag dal basalt formation, north greenland. rapport grønlands geologiske undersøgelse 99, 25–32. jepsen, h.f., escher, j.c., friderichsen, j.d. & higgins, a.k. 1994: the geology of the north-eastern corner of greenland – photogeological studies and 1993 field work. rapport grønlands geologiske undersøgelse 151, 21–33. jones, k.a. & escher, j.c. 1995: an e–w traverse across the caledonian fold belt from lambert land to norske øer. in: higgins, a.k. (ed.): express report: eastern north greenland and north-east greenland 1995, 23–41. unpublished report, geological survey of greenland, copenhagen. kalsbeek, f. & jepsen, h.f. 1983: the midsommersø dolerites geus bulletin 6.pmd 10-02-2005, 09:5455 56 and associated intrusions in the proterozoic platform of eastern north greenland – a study of the interaction between intrusive basic magma and sialic crust. journal of petrology 24, 605–634. kalsbeek, f. & jepsen, h.f. 1984: the late proterozoic zig-zag dal basalt formation of eastern north greenland. journal of petrology 25, 644–664. kalsbeek, f., nutman, a.p. & taylor, p.n. 1993: palaeoproterozoic basement province in the caledonian fold belt of northeast greenland. precambrian research 63, 163–178. kalsbeek, f., nutman, a.p., escher, j.c., friderichsen, j.d., hull, j.m., jones, k.a. & pedersen, s.a.s. 1999: geochronology of granitic and supracrustal rocks from the northern part of the east greenland caledonides: ion microprobe u-pb zircon ages. geology of greenland survey bulletin 184, 31–48. kalsbeek, f., thrane, k., nutman, a.p. & jepsen, h.f. 2000: late mesoproterozoic metasedimentary and granitic rocks in the kong oscar fjord region, east greenland caledonian fold belt: evidence for grenvillian orogenesis? journal of the geological society (london) 157, 1215–1225. kalsbeek, f., jepsen, h.f. & nutman, a.p. 2001a: from source migmatites to plutons: tracking the origin of c. 435 ma granites in the east greenland caledonian orogen. lithos 57, 1– 21. kalsbeek, f., jepsen, h.f. & jones, k.a. 2001b: geochemistry and petrogenesis of s-type granites in the east greenland caledonides. lithos 57, 91–109. leslie, a.g. & nutman, a.p. 2000: episodic tectono-thermal activity in the southern part of the east greenland caledonides. geology of greenland survey bulletin 186, 42–49. leslie, a.g. & nutman, a.p. 2003: evidence for neoproterozoic orogenesis and early high temperature scandian deformation events in the southern east greenland caledonides. geological magazine 140, 309–333. pedersen, s.a.s., craig, l.e., upton, b.g.j., rämö, o.t., jepsen, h.f. & kalsbeek, f. 2002: palaeoproterozoic (1740 ma) riftrelated volcanism in the hekla sund region, eastern north greenland: field occurrence, geochemistry and tectonic setting. precambrian research 114, 327–346. peel, j.s. 1980: geological reconnaissance in the caledonian foreland of eastern north greenland with comments on the centrum limestone. rapport grønlands geologiske undersøgelse 99, 61–72. peel, j.s. & smith, m.p. 1988: the wandel valley formation (early– middle ordovician) of north greenland and its correlatives. in: peel, j.s. (ed.): cambrian–jurassic fossils, trace fossils and stratigraphy from greenland. rapport grønlands geologiske undersøgelse 137, 61–92. rasmussen, j.a. & smith, m.p. 1996: lower palaeozoic carbonates in eastern north greenland, and the demise of the ‘sæfaxi elv nappe’. bulletin grønlands geologiske undersøgelse 172, 49–54. rasmussen, j.a. & smith, m.p. 2001: conodont geothermometry and tectonic overburden in the northernmost east greenland caledonides. geological magazine 138, 687–698. rejebian, v.a., harris, a.g. & huebner, j.s. 1987: conodont color and alteration. an index to regional metamorphism, contact metamorphism and hydrothermal alteration. bulletin geological society of america 99, 471–479. smith, m.p. 2000: cambro-ordovician stratigraphy of bjørnøya and north greenland: constraints on tectonic models for the arctic caledonides and the tertiary opening of the greenland sea. journal of the geological society (london) 157, 459–470. smith, m.p., sønderholm, m. & tull, s.j. 1989: the morris bugt group (middle ordovician – silurian) of north greenland and its correlatives. rapport grønlands geologiske undersøgelse 143, 5–20. smith, m.p., soper, n.j., higgins, a.k., rasmussen, j.a. & craig, l.e. 1999: palaeokarst systems in the neoproterozoic of eastern north greenland in relation to extensional tectonics on the laurentian margin. journal of the geological society (london) 156, 113–124. smith, m.p., higgins, a.k., soper, n.j. & sønderholm, m. 2004a: the neoproterozoic rivieradal group of kronprins christian land, eastern north greenland. in: higgins, a.k. & kalsbeek, f. (eds): east greenland caledonides: stratigraphy, structure and geochronology. geological survey of denmark and greenland bulletin 6, 29–39 (this volume). smith, m.p., rasmussen, j.a., robertson, s., higgins, a.k & leslie a.g. 2004b: lower palaeozoic stratigraphy of the east greenland caledonides. in: higgins, a.k. & kalsbeek, f. (eds): east greenland caledonides: stratigraphy, structure and geochronology. geological survey of denmark and greenland bulletin 6, 5–28 (this volume). sønderholm, m. & jepsen, h.f. 1991: proterozoic basins of north greenland. in: peel, j.s. & sønderholm, m. (eds): sedimentary basins of north greenland. bulletin grønlands geologiske undersøgelse 160, 49–69. steiger, r.h., hansen, b.t., schuler, c., bär, m.t. & henriksen, n. 1979: polyorogenic nature of the southern caledonian fold belt in east greenland. journal of geology 87, 475–495. strachan, r.a. & tribe, i.r. 1994: structure of the storstrømmen shear zone, eastern hertugen af orléans land, north-east greenland. in: higgins, a.k. (ed.): geology of north-east greenland. rapport grønlands geologiske undersøgelse 162, 103–112. watt, g.r., kinny, p.d. & friderichsen, j.d. 2000: u-pb geochronology of neoproterozoic and caledonian tectonothermal events in the east greenland caledonides. journal of the geological society (london) 157, 1031–1048. geus bulletin 6.pmd 10-02-2005, 09:5456 geological survey of denmark and greenland bulletin 7, 2004, p 41-44 recent research on danish groundwater has focused on clarifying the fate and transport of pesticides that leach through clayey till aquitards with low matrix permeability. previously, these aquitards were considered as protective layers against contamination of underlying groundwater aquifers due to their low permeability characteristics. however, geological heterogeneities such as fractures and macropores have been recognised as preferential flow paths within low permeable clayey till (e.g. beven & germann 1982). the flow velocities within these preferential flow paths can be orders of magnitude higher than in the surrounding clay matrix and pose a major risk of transport of contaminants to the underlying aquifers (e.g. nilsson et al. 2001). previous studies of transport in fractured clayey till have focused on fully saturated conditions (e.g. sidle et al. 1998; mckay et al. 1999). however, seasonal fluctuations of the groundwater table typically result in unsaturated conditions in the upper few metres of the clay deposits, resulting in different flow and transport conditions. only a few experiments have examined the influence of unsaturated conditions on flow and solute (the dissolved inorganic and organic constituents) transport in fractured clayey till. these include smallscale laboratory column experiments on undisturbed soil monoliths (e.g. jacobsen et al. 1997; jørgensen et al. 1998), intermediate scale lysimeters (e.g. fomsgaard et al. 2003) and field-scale tile drain experiments (e.g. kjær et al. 2005). the different approaches each have limitations in terms of characterising flow and transport in fractured media. laboratory studies of solute transport in soils (intact soil columns) are not exactly representative of field conditions due to variations in spatial variability and soil structure. in contrast, field studies hardly allow quantification of fluxes and mechanisms of transport. column and lysimeter experiments are often limited in size, and tile-drain experiments on field scale do not provide spatial resolution and often have large uncertainties in mass balance calculations. thus, in order to represent the overall natural fracture network systems on a field scale with respect to acquiring insights into flow and transport processes, the lysimeter needs to be larger than normal lysimeter size (< 1 m3). a modified large-scale lysimeter was therefore constructed by the geological survey of denmark and greenland (geus) at the avedøre experimental field site 15 km south of copenhagen (fig. 1). this lysimeter consisted of an isolated block (3.5 × 3.5 × 3.3 m) of unsaturated fractured clayey till with a volume sufficient to represent the overall preferential flow paths (natural fracture network) within lowpermeable clayey till at a field scale. 41 field experimental design for pesticide leaching – a modified large-scale lysimeter bertel nilsson, jens aamand, ole stig jacobsen and rené k. juhler geological survey of denmark and greenland bulletin 7, 41–44 (2005) © geus, 2005 avedøre denmark 50 km 0.00.0 0.5 1.5 2.5 3.5 4.5 1.0 2.0 3.0 4.0 depth (m) 0 50 100 horizontal fractures vertical fractures oxidised reduced caco3 free caco3 rich number of fractures per m2 r ed u ce d m at ri x z o n e 3 z o n e 2 z o n e 1 o x id is ed m at ri x 150 200 0.5 1.0 2.0 3.0 1.5 2.5 fig. 1. lithology and fracture frequency of the clayey till found at the avedøre field site where a large-scale lysimeter experiment was carried out. inset map: location of the field site. the brown colour indicates the distribution of the clayey till plain in denmark that was previously considered to be a protective layer against contamination of underlying aquifers from surface applications of pesticides. modified from mortensen et al. (2004). field site description the till plain at the avedøre field site is characterised by a 7 m thick, highly fractured clayey till aquitard covering a regional limestone aquifer. lithology, fracture systems and macropores have been measured at the study site (fig. 1) in order to describe the depositional environment (mckay et al. 1999). the fracture characterisation indicates that five distinct fracture systems are present in the till aquitard at the field site. two vertical fracture systems and one horizontal give the till a brick-like appearance between about 1.25 to 3.3 m depth (fig. 2). the till consists of massive and very stiff clayey material that has always caused problems to well borers and contractors in the copenhagen area due to the material hardness. it was therefore no surprise that difficulties were encountered during construction of tunnels below central parts of the city of copenhagen for the recently completed metro train system. modified large-scale lysimeter the lysimeter consists of an isolated till block (fig. 3), where the lower boundary is a steel plate (fig. 4), and the vertical walls around the block have been isolated with prefabricated bentonite plates to avoid water invasion from the surrounding environment. for monitoring and controlling transport through the block, four horizontal drainpipes were installed above the steel plate. the design of the modified lysimeter is described in detail in mortensen et al. (2004). 42 fig. 3. experimental set-up showing the large-scale lysimeter that consists of an isolated till block, and the adjacent monitoring and pumping wells. modified from mortensen et al. (2004). fig. 4. installation of the steel plate forming the base of the large-scale lysimeter. the steel plate was inserted by four hydraulic piston rods into the wall of the excavation at 3.3 m depth. fig. 2. clayey till as seen at 3.3 m depth at the avedøre field site. two vertical and one horizontal fracture systems give the till a brick-like appearance. the steel plate forms the bottom of the lysimeter. 3.5 m 3.5 m pumping well monitoring well till block infiltration area drainssampling groundwater level 3 .0 m 3 .3 m 0 .3 m fig. 5. computer-controlled spraying system providing controlled amounts of water to the till block to facilitate pesticide leaching. the pesticides and tracer compounds were added directly to the surface of the till block. the exposed upper surface of the block (infiltration area) is covered by a shelter to protect it from unmeasured contributions of rainfall. controlled artificial precipitation can be generated over the infiltration basin (fig. 5) using a computer-controlled spray system. the nozzles on the spray system are identical with nozzles traditionally used for agricultural pesticide spraying. pesticide leaching through an isolated block of clayey till a multiple tracer experiment and two different pesticideleaching experiments were carried out using the lysimeter, with precise control of the rain distribution and percolation through the lysimeter. the diffusive exchange of pesticides and multiple tracers between fractures and the matrix was examined in the lysimeter whereas specific sorption and degradation rates were determined in the laboratory. cored samples were collected in a 3.5 m deep excavation adjacent to the lysimeter for the laboratory work (fig. 6). results of the multiple tracer experiment are reported by mortensen et al. (2004). tracers during steady-state flow were transported quickly through the 3.3 m unsaturated clayey till block, with the first tracer being detected after about 25 minutes. multiple tracing techniques were applied to evaluate the importance of diffusive exchange on the overall transport processes. the main finding of the multiple tracer study was that there were large differences for the three different water fluxes used. the results of the pesticide leaching experiments provided some insights into the transport mechanisms in fractured clay, and verified a need for further work on leaching experiments in large-scale lysimeters. it is essential that transport mechanisms are addressed by the experimental conditions provided by lysimeters such as that described here. only through such experiments can plausible quantifications of mass fluxes be obtained. results of these studies have been reported in detail by aamand & jacobsen (2001), juhler & mortensen (2002), nilsson et al. (2002), mortensen et al. (2004) and aamand et al. (in press). 43 fig. 6. large excavation close to the lysimeter at the avedøre field site. the pit was excavated in steps with vertical and horizontal faces. profiles were orientated in two directions, perpendicular to one another, so that the fracture characterisation could be expressed in three dimensions. cored samples from matrix and fracture dominated parts were collected for degradation, sorption and pesticide diffusion studies in the laboratory. 44 acknowledgement this paper is an outcome of various combined laboratory and field studies (together named the avedøre project) initially carried out during 2000–2001 by the geological survey of denmark and greenland (geus) in collaboration with the technical university of denmark and københavns energi. a follow-up experiment was carried out during 2003–2004 by geus in collaboration with the geological institute, university of copenhagen. references aamand, j. & jacobsen, o.s. 2001: sorption and degradation of glyphosate and dichlobenil in fractured clay. in: walker, a. (ed.): pesticide behaviour in soils and water. bcpc symposium proceedings 78, 205–210. aamand, j., jacobsen, o.j. & nilsson, b. in press: glyphosats transport og omsætning i sprækket moræneler. danmarks og grønlands geologiske undersøgelse rapport. beven, k. & germann, m. 1982: macropores and water flow in soils. water resources research 18, 1311–1325. fomsgaard, i.s., spliid, n.h. & felding, g. 2003: leaching of pesticides through normal-tillage and low-tillage soil – a lysimeter study. i. isoproturon. journal of environmental science and health b38, 1–18. jacobsen, o.h., moldrup, p., larsen, c., konnerup, l. & petersen, l.w. 1997: particle transport in macropores of undisturbed soil columns. journal of hydrology 196, 185–203. jørgensen, p.r., mckay, l.d. & spliid, n.h. 1998: evaluation of chloride and pesticide transport in a fractured clayey till using large undisturbed columns and numerical analysis. water resources research 34, 539–553. juhler, r.k. & mortensen, a.p. 2002: analysing fluorobenzoate tracers in groundwater samples using liquid chromatography – tandem mass spectrometry – a tool for leaching studies and hydrology. journal of chromatography a957, 11–16. kjær, j., olsen, p., ullum, m. & grant, r. 2005: leaching of glyphosate and amino-methylphosphonic acid from danish agricultural field sites. journal of environmental quality 34, 608–621. mckay, l., fredericia, j., lenczewski, m., morthorst, j. & klint, k.e.s. 1999: spatial variability of contaminant transport in a fractured till, avedøre, denmark. nordic hydrology 30, 333–360. mortensen, a.p., jensen, k.h., nilsson, b. & juhler, r.k. 2004: multiple tracing experiments in unsaturated fractured clayey till. vadose zone journal 3, 633–644. nilsson, b., sidle, r.c., klint, k.e., bøggild, c.e. & broholm, k. 2001: mass transport and scale-dependent hydraulic tests in a heterogeneous glacial till-sandy aquifer system. journal of hydrology 243, 162–179. nilsson, b., aamand, j., jacobsen, o.s., juhler, r.k., mortensen, a.p. & broholm, m. 2002: udvalgte pesticiders transportveje og omsætning i sprækket moræneler i københavnsområdet. danmarks og grønlands geologiske undersøgelse rapport 2002/34, 62 pp. sidle, r.c., nilsson, b., hansen, m. & fredericia, j. 1998: spatially varying hydraulic and solute transport characteristics of a fractured till determined by field tests, funen, denmark. water resources research 34, 2515–2517. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: bn@geus.dk geological survey of denmark and greenland bulletin 7, 2004, p 73-77 73 dimension stones, or ornamental stones, are naturally occurring rocks that have properties that make them suitable for decorative exterior and interior use in the building industry. large parts of greenland should have a good potential for finding occurrences of valuable dimension stones, and the geological survey of denmark and greenland (geus) and greenland resources a/s (gras) therefore carried out prospecting in several parts of greenland in the summers of 2001–2004. the project is mainly financed by the government of greenland, but recently the european union and nuup kommunea have also contributed. dimension stone exploration in greenland all nordic countries with exposed crystalline basement have had large stone industries for decades, or even centuries. for example, the annual norwegian production of dimension stones and related materials had a value of about 100 million euros in 2003; even denmark has a well-established stone production and export from quarries on bornholm. greenland currently occupies an unfavourable position in this context, with only a small production of unpolished rock slabs, kerbstones and road material for local use; at present there is no export at all of natural stone products. however, a small company in nuuk (grønlandsk produktion a/s) has recently invested in facilities to process and polish natural stone, so with interest from local industry there is a good prospect for future dimension stone production. the development of a stone industry in greenland is compatible with the existing infrastructure for many reasons: (1) natural stone resources are plentiful and can be easily inspected because of the high degree of exposure in greenland, (2) the production only requires relatively low-cost investments, (3) areas of interest are often easily accessible from the sea, (4) the government of greenland favours new industrial initiatives, (5) the international dimension stone market demands a large variety of rock types, and special varieties – some of which are already known in greenland – are generally considerably more valuable than common rock types, and (6) the world market price for large blocks of special rock types is in the range of 800–1200 euros per cubic metre; with such price levels the financial costs of transportprospecting for dimension stones in greenland thomas v. rasmussen and hans kristian olsen geological survey of denmark and greenland bulletin 7, 73–76 (2005) © geus, 2005 ■ godthåbsfjord nuuk kangerluarsoruseq nassuttooq rifkol maniitsoq uummannaq tasiilaq 250 km narsaq kap farvel maarmorilik greenland sedimentary rockscrystalline rocks palaeogene basalts intrusive complexes caledonian orogenic belt palaeoproterozoic orogenic belts archaean craton cretaceous–palaeogene devonian–cretaceous lower palaeozoic middle–upper proterozoic ice 72°n 76°n 80°n 28°w36°w44°w52°w 68°n 64°n 60°n fig. 1. simplified geological map of greenland showing place names and areas prospected for dimension stones. the gardar province of south greenland, including the eriksfjord formation and the ilímaussaq alkaline complex, is situated between narsaq and the inland ice. 74 ing the blocks from greenland to worldwide markets only present a minor problem. there is currently a trend of depopulation of villages in greenland due to a shortage of local economic opportunities. the potential benefits of establishing dimension stone quarries near local communities would help sustain the existing smaller villages in greenland, which has a high priority on the political agenda. the dimension stone industry is environmentally friendly and labour intensive, and the value of the mined product can be high, making exploitation of the natural resource both socially and environmentally beneficial. previous dimension stone activities in greenland greenland previously hosted a commercial quarry at maarmorilik in central west greenland (fig. 1), where marble from the palaeoproterozoic mârmorilik formation was excavated and exported to denmark between 1968 and 1972. several well-known public and private buildings in denmark have floors or facades of this greenlandic marble. the igaliko sandstone from the mesoproterozoic eriksfjord formation in south greenland (fig. 1) has been exploited for local use as building stones, e.g. in the settlement of igaliko near narsaq. since 1974, however, and up to 2000 very little effort has been directed towards prospecting for dimension stone in greenland. the exception was a reconnaissance study in 1990, covering the west coast of greenland between uummannaq and kap farvel (garde et al. 1991; gothenborg et al. 1994). criteria for suitable dimension stone localities a good dimension stone locality must contain an unusual and attractive rock type of high technical quality, which can fetch a high price on the international market. the fracture density should be low, with no more than three sets of fractures that must be perpendicular to each other, so it is possible to produce blocks of about 1 × 2 × 3 m in size. furthermore, the locality must contain a large volume of homogeneous rock that can easily be quarried without removal of excessive overburden or quarrying below sea level, and without hazards such as danger from falling blocks. to minimise the cost of local transport, the site must be accessible from the sea. in spite of these strict criteria greenland offers a large potential, not least in view of its varied bedrock geology, suitable topography, excellent exposure and very long accessible coastline. prospected areas in 2001–2004 the target areas for the surveys in 2001–2004 were chosen with background in the previous reports on dimension stones and the general criteria listed above, the geological literature, and field observations by geologists from the survey in recent years. the level of the existing local infrastructure was also taken into account. the main target areas were located in the nassuttooq (nordre strømfjord) and nuuk regions in central west greenland, the gardar province in south greenland and around tasilaq in east greenland (fig. 1). in each area, potential dimension stone localities were mapped and very large samples were collected for testing, cutting and polishing. a b c fig. 2. a–c: polished slabs of migmatitic biotite gneisses from nassuttooq. width of slabs 30 cm. central west greenland in the central and inner parts of nassuttooq the primary targets were occurrences of large areas of migmatitic biotite gneiss and orthopyroxene gneiss (charnockite) with low fracture density (rasmussen 2003; rasmussen & olsen 2003). about ten such localities of migmatitic biotite gneiss with low fracture density were identified in the nassuttooq region. the migmatitic biotite gneiss has spectacular textures with pinkish red and grey colours (fig. 2a–c). furthermore, several varieties of orthopyroxene gneiss (fig. 3a, b) with low fracture density were located in the nassuttooq region and on the islands rifkol and maniitsoq just off the coast north of nassuttooq. nuuk–kangerluarsoruseq region in the region around nuuk, the capital of greenland, and nearby areas to the south several interesting rock types were located. in kangerluarsoruseq (færingehavn) south of nuuk a massive olivine dolerite/gabbro and an anorthosite were sampled, as well as a leucogabbro with attractive iridescence in the feldspars (peristerite and labradorite). closer to nuuk there are several localities with augen gneiss and granite, both with low fracture densities. just east of nuuk on the island of sermitsiaq there is a potential for dark green, brown and black ultramafic metamorphic rocks. south greenland in south greenland the main focus of the 2001–2004 surveys was on the ilímaussaq alkaline complex and the sandstones of the eriksfjord formation (kalvig et al. 2002; rasmussen 2003; rasmussen & olsen 2003). the ilímaussaq alkaline complex has the advantage that several unique and attractive rock types with very spectacular colours are exposed within a small area. two types of kakortokite (fig. 3c) and naujaite (fig. 3d) were sampled. the main challenge of the ilímaussaq alkaline complex is to locate areas with a sufficiently low fracture density, as many exposures are heavily fractured in the surface outcrops. the eriksfjord formation close to the ilímaussaq alkaline complex hosts several occurrences of colourful sandstones. the common red type (fig. 4a) has a good potential because many old buildings in europe built of red sandstone are in need of restoration, and suitable sandstone resources in europe are gradually becoming exhausted. the sandstones from the eriksfjord formation are quartz-cemented, which is a distinct advantage compared to the more common calcitecemented sandstones. quartz-cemented sandstones can be used in all environments because they are resistant to chemia b c d fig. 3. a: polished slab of homogeneous orthopyroxene gneiss from nassuttooq. b: polished slab of garnet-biotite-quartz paragneiss from nassuttooq. c: polished slab of kakortokite from the ilímaussaq alkaline complex of south greenland. d: polished slab of naujaite from the ilímaussaq alkaline complex of south greenland. width of slabs 20 cm. 75 76 cal weathering, and also takes polishing very well. the red sandstones of the eriksfjord formation are intruded locally by gardar dolerite dykes, which have baked the sandstone close to the dykes and transformed it into a massive quartzite with an attractive blue-green colour (fig. 4b). although the volume of this lithological variety is limited, it has promising possibilities on the international dimension stone market. east greenland in the ammassalik area of east greenland large outcrops of massive pink granite with extremely low fracture density have been located (kalvig et al. 2002; rasmussen 2003). current and future dimension stone projects the next steps for dimension stone prospecting in greenland will be promotion of the collected rocks to large international stone companies and continued field prospecting in greenland. in 2004, the dimension stone project was granted 50 000 euros from the eu’s northern periphery programme (npp). under the npp the nordic countries (sweden, norway, finland and denmark) have formed a working group (pnastina), which aims at promoting the natural stone industry in northern europe. since large international stone companies are already established in finland, norway and sweden, the pnastina group provides an important contact for denmark and greenland to the international dimension stone market. the plan for 2005 is to continue the prospecting in the godthåbsfjord region of southern west greenland and in the gardar province in south greenland. acknowledgements the bureau of minerals and petroleum (bmp), nuup kommunea, sulisa a/s and the eu (northern periphery programme) provided financial support for the project. references garde, a.a., bugnon, c. & gothenborg, j. 1991: ornamental stones in west and south greenland. rapport grønlands geologiske undersøgelse 152, 50–55. gothenborg, j., garde, a.a. & bugnon, c. 1994: greenland ornamental stone resources. the 1990/91 ornamental stone project. open file series grønlands geologiske undersøgelse 94/2, 143 pp. kalvig, p., knudsen, c.n. & rasmussen, t.v. 2002: potentialer for facadesten og skærver i grønland. danmarks og grønlands geologiske undersøgelse rapport 2002/11, 104 pp. rasmussen, t.v. 2003: dimension stone prospecting in west, south and east greenland 2002. danmarks og grønlands geologiske undersøgelse rapport 2003/8, 60 pp. rasmussen, t.v. & olsen, h.k. 2003: dimension stone prospecting in west and south greenland 2003. danmarks og grønlands geologiske undersøgelse rapport 2003/107, 67 pp. a b fig. 4. a: polished slab of red sandstone from the eriksfjord formation in south greenland. b: polished slab of massive blue-green quartzite from the eriksfjord formation in south greenland. width of slabs 20 cm. authors’ addresses t.v.r., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: tvr@geus.dk h.k.o., greenland resources. present address: nunaoil a/s, tuapannguit 38, p.o. box 579, dk-3900 nuuk, greenland. geological survey of denmark and greenland bulletin 34, 2015, 1-86 geological survey of denmark and greenland bulletin 34· 2015 the ammonites of the middle jurassic cranocephalites beds of east greenland john h. callomon, peter alsen & finn surlyk with an appendix by peter alsen: description of a new bajocian (middle jurassic) ammonite species, cranocephalites tvaerdalensis sp.nov., from geographical society ø, north-east greenland geological survey of denmark and greenland ministry of energy, utilities and climate geological survey of denmark and greenland bulletin 34 keywords biostratigraphy, middle jurassic, jameson land basin, boreal realm, bajocian, borealis zone, indistinctus zone, pompeckji zone, intermissus subzone, carlsbergensis subzone, gracilis subzone, episcopalis subzone, cranocephalites carolae sp. nov., cranocephalites intermissus sp. nov., cranocephalites episcopalis sp. nov. cover illustration jurassic sedimentary strata exposed on the south-western side of the upper ugleelv valley in eastern jameson land. such exposures along ugleelv yielded rich middle jurassic ammonite faunas that formed the key to the detailed taxonomic and chronostratigraphic results presented in this bulletin. the height of the plateau above the valley floor is about 500 m. photo: michael engkilde. frontispiece: facing page john h. callomon in the ‘greenland collections’ room, at that time on the second floor of the geological museum (university of copenhagen), repository of the material studied in this work. photo: peter alsen. chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, department of geoscience, aarhus university; minik rosing, geological museum, university of copenhagen; finn surlyk, department of geosciences and natural resource management, university of copenhagen scientific editor of this volume: jon r. ineson editorial secretary: jane holst referees: eckhart mönning (germany) and paul l. smith (canada) illustrations: jette halsskov digital photographic work: benny m. schark layout and graphic production: henrik klinge pedersen printers: rosendahls · schultz grafisk a/s, albertslund, denmark manuscript received: 30 april 2014 final version approved: 19 october 2015 printed: 28 december 2015 issn (print) 1604-8156 issn (online) 1904-4666 isbn (print) 978-87-7871-427-5 isbn (online) 978-87-7871-428-2 citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bulletin 34, 145 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk or at www.geus.dk/publications/bull „ de nationale geologiske undersøgelser for danmark og grønland (geus), 2015 for the full text of the geus copyright clause, please refer to www.geus.dk/publications/bull 4 contents abstract ..................................................................................................................................................................................... 7 preface ...................................................................................................................................................................................... 8 introduction .......................................................................................................................................................................... 9 history of research ......................................................................................................................................................... 9 early days ................................................................................................................................................................... 9 the koch years, 1926–1951: initial geological exploration ...................................................................... 9 1957–1993: detailed biochronology ................................................................................................................ 10 1994–1996: ugleelv, central jameson land ................................................................................................. 19 stratigraphy .......................................................................................................................................................................... 21 localities .......................................................................................................................................................................... 21 lithostratigraphy .......................................................................................................................................................... 21 jameson land ......................................................................................................................................................... 22 ugleelv – northern hurry inlet ........................................................................................................................ 25 biostratigraphy .............................................................................................................................................................. 45 standard zone stratigraphy.................................................................................................................................. 45 taxonomy: general principles ...................................................................................................................................... 48 systematic taxonomy ....................................................................................................................................................... 50 explanations and abbreviations ........................................................................................................................ 50 order ammonoidea .................................................................................................................................................... 51 suborder ammonitina ................................................................................................................................................ 51 superfamily stephanoceratoidea neumayr 1875 ........................................................................................ 51 family cardioceratidae siemiradzki 1891 .................................................................................................... 51 subfamily arctocephalitinae meledina 1968 ............................................................................................... 51 genus cranocephalites spath 1932 ........................................................................................................... 51 borealis standard zone .............................................................................................................................................. 52 bo-1: cranocephalites borealis (spath 1932) sensu stricto, trans α ............................................. 53 bo-2: cranocephalites borealis (spath 1932) trans β ..................................................................... 54 bo-3: cranocephalites borealis (spath 1932) trans γ ..................................................................... 55 general discussion of the cranocephalites borealis group ............................................................ 56 indistinctus standard zone ....................................................................................................................................... 57 in-1, in-2: cranocephalites spp. (u-1), (u-2) ................................................................................. 57 in-3 – in-8: cranocephalites indistinctus callomon 1959 ................................................................... 58 in-3: cranocephalites indistinctus trans α (sensu callomon 1993) ........................................... 58 in-4: cranocephalites indistinctus trans α´......................................................................................... 58 in-5: cranocephalites ex. gr. indistinctus (u-3) ............................................................................... 59 in-6: cranocephalites indistinctus callomon 1993 sensu stricto, trans β .................................. 59 in-7: cranocephalites ex. gr. indistinctus (u-4) ............................................................................... 59 in-8: cranocephalites indistinctus trans γ ......................................................................................... 60 pompeckji standard zone ......................................................................................................................................... 60 intermissus subzone ............................................................................................................................................. 60 po-1: cranocephalites carolae sp. nov. sensu stricto, trans α ......................................................... 60 po-2: cranocephalites carolae sp. nov. trans β .................................................................................. 61 po-3: cranocephalites intermissus sp. nov. trans α .......................................................................... 62 po-4: cranocephalites intermissus sp. nov. sensu stricto, trans β .................................................. 62 po-5: cranocephalites sp. (u-5) ........................................................................................................... 63 po-6: cranocephalites sp. aff. intermissus (u-6) .............................................................................. 63 po-7: cranocephalites sp. aff. intermissus (u-7) .............................................................................. 63 carlsbergensis subzone ...................................................................................................................................... 67 po-8 – po-9: the group of cranocephalites pompeckji madsen – furcatus spath ........................ 67 po-8: cranocephalites pompeckji (madsen 1904) ........................................................................... 67 po-9: cranocephalites furcatus spath 1932 ...................................................................................... 69 po-10 – po-14: the group of cranocephalites carlsbergensis callomon ......................................... 70 po-10 – po-12: cranocephalites carlsbergensis callomon .................................................................... 70 5 po-10: cranocephalites carlsbergensis trans α ..................................................................................... 71 po-11: cranocephalites carlsbergensis sensu stricto, trans β ............................................................ 71 po-12: cranocephalites carlsbergensis trans γ ..................................................................................... 74 po-13 – po-14 faunal horizons ................................................................................................................... 74 po-13: cranocephalites tvaerdalensis alsen 2015.............................................................................. 74 po-14: cranocephalites sp. aff. tvaerdalensis (u-8)........................................................................... 75 gracilis subzone ..................................................................................................................................................... 76 po-15: cranocephalites gracilis spath 1932 ....................................................................................... 76 po-16: cranocephalites sp. aff. gracilis spath 1932 (u-9) ............................................................ 78 po-17: cranocephalites transitorius (spath 1932) trans α .............................................................. 78 po-18: cranocephalites transitorius (spath 1932) sensu stricto, trans β ...................................... 78 po-19: cranocephalites maculatus spath 1932.................................................................................. 80 po-20: cranocephalites ornatus (spath 1932).................................................................................... 80 episcopalis subzone .............................................................................................................................................. 80 po-21: cranocephalites episcopalis sp. nov. sensu stricto, trans α ................................................... 80 po-22: cranocephalites episcopalis sp. nov. trans β ........................................................................... 81 po-23: cranocephalites sp. (u-10) ........................................................................................................ 81 conclusions .......................................................................................................................................................................... 82 acknowledgements ........................................................................................................................................................... 82 references ............................................................................................................................................................................. 83 plates 1 –20 ........................................................................................................................................................................... 87 appendix 1 (including 5 plates) ................................................................................................................................. 129 6 7 abstract authors’ addresses j.h.c. (deceased), university college london, gower street, wc1e 6bt, london, uk. p.a. (corresponding author), geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen, denmark. e-mail: pal@geus.dk f.s., department of geosciences and natural resource management, university of copenhagen, øster voldgade 10, dk1350 copenhagen k, denmark. callomon, j.h., alsen, p. & surlyk, f. 2015: the ammonites of the middle jurassic cranocephalites beds of east greenland. geological survey of denmark and greenland bulletin 34, 145 pp. thick successions of marine middle jurassic deposits rich in ammonites occur in the jameson land basin in central east greenland. the fauna of the so-called cranocephalites beds of this basin, comprising the borealis–pompeckji standard zones, was until now largely represented by a single collection. this was made by t.m. harris during a 1927 excursion up the valley of ugleelv to katedralen, the type area of cranocephalites pompeckji, which is the oldest named species of this genus. revisits to this area in 1994 and 1996 by jhc resulted in a large bed-by-bed collection of cranocephalites. the number of faunal horizons that could be distinguished grew from the nine previously recognised to thirty-four today. the zonal stratigraphy of the cranocephalites beds encompasses the borealis, indistinctus and pompeckji standard zones. the pompeckji zone is subdivided into four new subzones, reflecting four successive basic morphologies of cranocephalites that should be recognisable more widely and are thus useful for subzonal correlations. the detailed zonation that serves as the secondary standard zonation for the boreal province in the middle jurassic is thus highly improved. the biostratigraphic resolution obtained here is near the achievable limits. it allows a high-resolution study of the evolution of the ammonites which on this timescale appears to be continuous. three new species are described: cranocephalites carolae sp. nov., cranocephalites intermissus sp. nov. and cranocephalites episcopalis sp. nov. an additional new species, cranocephalites tvaerdalensis sp. nov., is described in the appendix by p. alsen based on collections from tværdal on geographical society ø, north-east greenland. this species is also recorded in jameson land. 8 professor john h. callomon died on 1 april 2010. for many years, he had been occupied with resolving the faunal succession of the middle jurassic cranocephalites beds in east greenland. when he died he left an unfinished manuscript, which his family entrusted to peter alsen and finn surlyk to complete. john h. callomon had an interest in geology since childhood, but in his professional career he was a distinguished professor in his field of inorganic chemistry and wrote landmark papers in spectroscopy. his passion for jurassic ammonite stratigraphy developed during his student years in oxford where he also met desmond t. donovan, through whom john was first introduced to east greenland as a member of an expedition led by lauge koch. he thought he was to assist donovan during the 1957 field season, but after a few days donovan flew off to wollaston forland and john was left with the responsibility to collect and record the upper jurassic sections in milne land. he subsequently participated in a number of expeditions to east greenland and collaborated closely with the late tove birkelund and fs. the outcome was a number of papers on middle and upper jurassic ammonite stratigraphy. in 1993, he was awarded the prestigious steno medal by the danish geo logical society in recognition of the influence of his work for the understanding of the jurassic in east greenland. preface field camp in the uglelev valley in the summer of 1971. from left to right: john callomon, tove birkelund, lise alkjærsig (field assistant) and finn surlyk. 9 jurassic sedimentary rocks are widely distributed in central east greenland and north-east greenland from jameson land in the south to store koldewey in the north (fig. 1a). they are commonly beautifully exposed and a comprehensive review was given by surlyk (2003). lower jurassic rocks are restricted to the southern part of the east greenland rift complex, in the jameson land basin (fig. 1b). the kap stewart group comprises lacustrine and fluvial deposits of rhaetian–sinemurian age overlain by the shallow marine neill klinter group of pliensbachian – early bajocian age (dam & surlyk 1998). jurassic rifting was initiated in the late bajocian and a large area north of the jameson land basin was transgressed for the first time since the late permian. major n–s-trending faults, which sidestepped en echelon to the east, controlled the margins of the jurassic rift complex. the main sedimentary influx took place in the relay-ramp areas between the faults and the rift thus comprised several marine gulfs with their heads towards the north; the main sedimentary transport direction was axial towards the south. rifting intensified during the jurassic to reach a climax in mid-volgian time (surlyk 2003). the key area for the study of middle jurassic bio stratigraphy is jameson land where thick successions, commonly rich in ammonites, occur (fig. 1b). exposures along hurry inlet and around the river of ugleelv are particularly important in this respect and this is where the ammonite faunas described in the classical monograph of spath (1932) were collected. other important, but less accessible areas occur in central and northern jameson land; all the localities are shown on the maps in figs 2–7. the upper bajocian – middle callovian part of the early rift succession in jameson land is placed in the pelion and fossilbjerget formations, as described below. the jurassic in east and north-east greenland has served as an important analogue for contemporaneous successions in the northern north sea and along the conjugate norwegian margin. in addition, the ammonite faunal succession in the middle jurassic is the most complete in the arctic and the zonation serves as a standard of reference for the boreal province (callomon 1993). history of research the history of the knowledge of the middle jurassic ammonites of east greenland can be resolved into four phases. early days the first middle jurassic ammonites from east greenland were collected by nikolaj hartz, henrik deichmann and otto nordenskjöld just over a century ago, during the danish expedition in 1900 led by georg carl amdrup and hartz. the collections were described by madsen (1904), who correctly assigned their ages as middle jurassic, probably bathonian–callovian, although he could not be more precise. he also relayed reports of “innumerable ammonites scattered about on the plateaux inland”, that is, in jameson land west of the northern reaches of hurry inlet (madsen 1904, p. 169). this expedition, like others of this era, had been one of general natural science and geographic exploration. the koch years, 1926–1957: initial geological exploration the systematic study of the geology of east greenland began with the many expeditions led by lauge koch in the years 1926–1958. the widespread presence of richly fossiliferous middle jurassic marine sediments became apparent during the first of the expeditions in 1926–1927, when large collections of ammonites were made by alfred rosenkrantz and tom harris in jameson land west of hurry inlet in central east greenland. they were described in a classical monograph by spath (1932), which for many years remained the primary reference as it still is today, at least as concerns the description of species. what was lacking was a detailed stratigraphical framework. this was largely inevitable as the geological observations inland had to be made in the spring on sledge journeys over still largely snow-covered ground. the relative stratigraphic successions of forms found at widely scattered localities had in part to be inferred and the type horizons of most of them were not precisely known. all in all, the succession of ammonites seemed to characterise a series of four bio stratigraphical units: kepplerites–cadoceras beds (highest) arcticoceras beds arctocephalites beds cranocephalites beds (lowest) callomon and earlier authors informally used the term ‘beds’, which in modern lithostratigraphical nomenclature introduction 10 refers to formal subdivisions of members; the term is accordingly given in lower case here. these four units provide an overall frame of reference that continues to be useful today. the lowest unit rests on a thick succession of dark mudstones with only scattered fossils, and no ammonites, the sortehat formation of dam & surlyk (1998) and surlyk (2003), now referred to the neill klinter group (dam & surlyk 1998). within these four middle jurassic biostratigraphic units, it seemed possible to distinguish some five ammonite faunal horizons, with indications of perhaps two more – a total of seven (spath 1932). another rich collection from the middle jurassic of the coastal cliffs of neill klinter, along the western shores of hurry inlet, was made in 1933 by hermann aldinger, who also produced excellent geological maps and measured the first usable sections (aldinger 1935). his collections, although also sent to london for analysis by spath, were however never described. his published account of the succession was fully compatible with the biostratigraphy of spath (1932) but added nothing of biostratigraphic significance. further extensive collections were made in central jameson land by stauber (1940) during horseback traverses in 1938 but these, too, were left undescribed; the collections were almost wholly devoid of stratigraphical information. after the second world war, exploration was resumed north of kong oscar fjord (72°n), on the islands of traill ø and geographical society ø (fig. 1a), but although the middle jurassic there is thickly developed and extensive, fossils, including ammonites, are scarce and scattered. two further ammonite horizons were, however, recognised by donovan (1955), bringing the total up to perhaps nine. the state of knowledge at the time was comprehensively reviewed by donovan (1957). 1957–1993: detailed biochronology a new phase in the study of the middle jurassic started with koch’s last two expeditions in 1957 and 1958. it was prompted by a re-examination of some serious unresolved problems posed by the faunas already described, relating to their ages. these problems are reflected in the names given by spath (1932) to three of his four biostratigraphical units cited above. the highest of these was named after the genera kepplerites, of the family kosmoceratidae, and cadoceras, of the family cardioceratidae. these families are well represented in the classical jurassic succession of europe, where as far as was then known, kepplerites and cadoceras made a sudden appearance somewhere in the lower but not lowest part of the callovian. the greenland forms, while sim18°w 16°w 74°n 76°n 72°n 22°w 20°w cretaceous jurassic triassic permian fault buried deepseated faults greenland milne land traill ø geographical society ø hold with hope kuhn ø hochstetter forland store koldewey wollaston forland jameson land kong oscar fjord hurry inlet liverpool land a 100 km 26°w 22°w24°w28°w 26°w 24°w fig. 1. a: simplified geological map showing the distribution of per mian–cretaceous sediments in east and north-east greenland and with the study area of jameson land indicated. b (facing page): map of jameson land, showing the outcrops of the middle jurassic and the locations of the sections that had been recorded up to 1993, numbered 1–86 from south to north (callomon 1993). section numbers are indicated on selected sections; for more detailed positioning of sections see figs 2–7. 11 86 82 fig. 7 fig. 2 fig. 3 fig. 4 fig. 6 fig. 5 79 81 59 78 73 69 62 55 52 51 7467 60 56 58 48 46 40 38 41 35 27 25 21 19 24 13 12 10 6 1 3230 49 2850 29 84 ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ 24°w middle jurassic outcrops sections of callomon (1993) mountain 22°w 72°n 71°n antarctic havn flem ing fjo rd carls ber g f jord kong oscar fjord sc hu ch er t d al ørste d dal scoresby land jameson land liverpool land scoresby sund h ur ry in let 25 km olym pen pothorst bjerg e pelio n kosmocera s bjerg parnas depot elv kated rale n sorteh at spath fje ld vardeklø ft mikae l bjerg langryg gen fossilb jerg et tref jord bjerg hareelv ugleelv b 12 70°45'n 70°40'n 70°30'n 5 km 22°40'w22°50'w prim ula elv 9 8 7 hareelv astartekløft moskusoksekløft skævdal dinosaurus kløft la ks ee lv os tre ae lv m øn se lv constable pynt 500 zackenberg 12 11 10 781 800 fortet harris fjeld kap stewart 150 230 230 387 460 550 freyberg fjeld 670 1 2 3 4 5 6 770 814 brinkmann fjeld spath fjeld 720 j. p. koch fjeld 909 h ur ry in le t varde kløft goniomyakløft sill dyke altitude (metres above sea level) ice legend (figs 2–7, 9) recent and quaternary drift recent alluvial fans sortehat fm (n.k. gp) pelion fm fossilbjerget fm neill klinter gp (excl. sortehat fm) fleming fjord fm kap stewart gp raukelv fm hareelv fm olympen fm hesteelv fm lower cretaceous upper jurassic middle jurassic triassic rhaetian – lower jurassic upper jurassic – lower cretaceous pingodal fm gipsdalen fm wordie creek fm extrusives syenite intrusion n ei ll k lin te r fig. 2. geological map showing locations of the sections 1–12 recorded by callomon (1993) along neill klinter on the western shores of hurry inlet. map section from geological map sheets 70 ø.1 syd (bengaard et al. 1986) and 70 ø.1 nord (friderichsen & surlyk 1981). the accompanying legend is applicable to figs 2–7 and fig. 9. n.k. gp: neill klinter group. 13 ?? ? ? ? 71°05'n 71°00'n 70°55'n 70°50'n 70°45'n 23°00'w 22°50'w 22°40'w 5 km 25 26 27 23 22 21 20 19 16 18 17 13 1415 24 550 640 561 107 610 640 610623 540 langryggen sortedal elis teebjerg katedralen sortehat ugleelv ræ ve elv fa lke elv gåseelv dusén k litdal bjerg bjerg h ur ry i nl et n a t h o r s t f j e l d fig. 3. geological map of the area around ugleelv with the names of places and locations of sections 13–27 recorded by callomon (1993). map section from geological map sheets 70 ø.1 nord (friderichsen & surlyk 1981) and 71 ø.1 syd (birkelund & higgins 1980). for legend, see fig. 2; red box shows area of fig. 9. 71°25'n 71°20'n 71°15'n 71°10'n 22°30'w23°00'w 22°50'w 22°40'w 5 km5 km 48 47 46 38 39a 39b 40 45 44 43 42 41 36 37 34 35 33 32 fossilbjerget mikael bjerg gule horn liasryggen hjø rn ef je ld et depotelv trefjord bjerg 751 730 960 880 950 1050 1060 880 790 900 910 622 630 660 690 centralbjerg skansen le pid op ter isel v passage n k lit da l c ar ls be rg f jo rd lejrelv lias elv 840 751 passage n fig. 4. geological map of the area around fossilbjerget and trefjord bjerg, showing locations of sections 32–48 recorded by callomon (1993). map section from geological map sheet 71 ø.1 syd (birkelund & higgins 1980). for legend, see fig. 2. 15 ilar, were, however, not conspecific with the european forms. nevertheless, it seemed safe to assign a callovian age to them also, making them slightly older than the oldest european forms but squeezing them into the small bio stratigraphical gap still left in the lower callovian below their first appearance there (spath 1932). the ammonites from the three units below the cadoceras–kepplerites beds were, however, totally unknown in europe. this is reflected by the new generic names they were given: cranocephalites, arctocephalites and arcticoceras, in ascending order. they had, however, been previously recorded from elsewhere in the arctic, for example in petshora (keyserling 1846, sokolov 1912), franz josef land (newton & teall 1897; whitfield 1907) and novaya zemlya (sokolov 1913; salfeld & frebold 1924). the evidence from east greenland therefore suggested the existence of four successive ammonite associations, of which only the highest closely resembled european lower callovian forms (spath 1932). while unknown in the underlying european bathonian, the three lower assemblages did, however, have a certain resemblance to the forms of another european family, the macrocephalitidae, which dominate the lowest callovian. the most cautious approach seemed therefore to assume that the morphological differences between the greenland and european forms were perhaps a reflection of bioprovincial differentiation, making arctocephalites, cranocephalites and macrocephalites but races of the macrocephalitidae from the lower callovian, or perhaps uppermost upper bathonian. the corollary was that there were possibly no bathonian ammonites in greenland, nor, by correlation, anywhere else in the arctic. an implicit assumption that hence there were also no sediments of bathonian age in the arctic was not well founded. very similar interpretations had meanwhile also been applied to closely analogous – and equally fallacious – observations on the middle jurassic ammonites of the whole of the american cordillera, from the andes to alaska. these, too, could be interpreted as bioprovincially differentiated relatives of the macrocephalitidae. together, these interpretations did much to cause arkell, in his influential book (arkell 1956) to illustrate a more general picture of the bathonian as a time of world-wide marine sea-level fall leaving no sediments. attempts to date these arctic faunas more securely could make no progress without new evidence. the efforts of callomon were therefore concentrated in 1957–1958 on the ammonite biostratigraphy of southern jameson land, making new in-place collections from carefully recorded sections. it quickly emerged that an immensely rich succession of faunas was to be discovered. after four weeks in the field, the number of faunal horizons characterised by clearly distinguishable ammonite assemblages had risen to nine (callomon 1959). the resumption of systematic mapping of the sedimentary basins in central east greenland by the university of copenhagen in the years 1968–1974 provided access to the remoter interior of jameson land, where a wealth of localities and ammonite assemblages was discovered (birkelund et al. 1971). in all, over 80 sections yielding middle jurassic ammo nites were recorded (figs 1–7). the locations of the more important ones relevant to the present discussions are shown in figs 2–4. the number of well-characterised, time-ordered, stratigraphically precisely located faunal horizons was now raised to thirty-seven. this biostratigraphical framework formed the basis of a standard stratigraphical classification down to zonal level. callomon (1995) presented his views on the relationships between bioand chronostratigraphy which are at variance with generally accepted distinctions between the two concepts as he saw his so-called standard zones, which are essentially chronozones, as basic units in chronostratigraphy. the stratigraphic results were reviewed by callomon (1993, 2003) and are summarised in fig. 8. the vertical timescale is drawn on an equispaced faunal horizon approximation, the horizons recognised by callomon (1993) being numbered consecutively upwards as shown on the right in the figure. the species of kepplerites found in the middle of the succession, at horizon 25, is so similar to that found in the kepplerites keppleri horizon defining the base of the callovian in europe (callomon & dietl 2000) that the bathonian age of the underlying 24 horizons could no longer be in doubt. there were in fact strong grounds, based on a chain of correlations via the east pacific, for believing that the lowermost faunal horizon in greenland, that of cranocephalites borealis, was of early late bajocian age (callomon 1985a). such a correlation has meanwhile received support from strontium stable-isotope stratigraphy based on belemnites (m. engkilde, personal communication 1998). it had become clear that the upper bajocian and bathonian ammonite succession of central east greenland was now the best known and nearest to complete anywhere within the clearly delimited boreal province, whose ammonites were so strongly endemically confined that precise correlations with their temporal equivalents elsewhere in the world remained impossible: there were no known areas of overlap. a separate bioprovincial secondary standard chronozonation (in the sense of callomon 1985b) for the boreal province was therefore to be constructed in parallel with the primary standard zonation (fig. 8). the primary standard zonation has also been drawn on a european equispaced faunal horizon approximation by callomon & chandler (1990) for the bajocian, westermann & callomon 16 co lo ra do da l 23 °1 0'w 23 °2 0'w 23 °3 0'w 23 °4 0'w 23 °5 0'w 24 °0 0'w 24 °1 0'w 71 °2 5'n 71 °3 0'n 71 °3 5'n 83 1 94 0 88 6 82 0 10 85 11 12 12 81 10 05 63 4 65 8 65 3 81 5 12 45 10 75 10 29 58 4 80 81 78 77 76 75 74 72 68 71 70 69 65 64 63 62 61 55 56 57 58 pe lio n o ly m pe n ko sm oc er as b je rg potho rst bjerg e pa rn as fegin elv jur ael v lo din el v 60 66 59 a 59 b 67 73 79 5 km pingel dal re gn ee lv jen s m un k pl at ea u fi g. 5 . g eo lo gi ca l m ap o f t he ar ea ar ou nd th e m ou nt ain s o ly m pe n, p ar na s a nd p eli on sh ow in g lo ca tio ns o f s ec tio ns 5 5– 81 re co rd ed b y c all om on (1 99 3) . m ap se ct io n fro m g eo lo gi ca l m ap sh ee ts 71 ø .1 s yd (b irk elu nd & h ig gi ns 1 98 0) , 7 1 ø .2 s yd (f rid er ich se n & b ro m ley 1 97 6) , 7 1 ø .2 n or d (h en rik se n & p er ch -n iel se n 19 77 ) a nd 7 1 ø .1 n or d (p er ch -n iel se n et al . 1 98 3) . f or le ge nd , s ee f ig . 2 . 17 36 3 46 7 30 5 48 6 21 6 18 2 11 9 13 0 13 4 17 3 42 1 28 6 56 9 49 1 75 1 58 3 23 7 5 km 23 °1 0'w71 °2 0'n 71 °1 5'n 71 °1 0'n 23 °3 0'w 23 °5 0'w 24 °0 0'w 32 2 54 53 52 51 50 31 30 49 29 28 jac ob s ev er in bj er g pingel d al de po tel v olym pelve n jur ael v feg in e lv dogg ere lv m ik ae l b je rg fi g. 6 . g eo lo gi ca l m ap o f t he ar ea ar ou nd m ik ae l b jer g s ho wi ng lo ca tio ns o f s ec tio ns 2 8– 31 an d 49 –5 4 re co rd ed b y c all om on (1 99 3) . m ap se ct io n fro m ge ol og ica l m ap sh ee ts 71 ø .1 s yd (b irk elu nd & h ig gi ns 19 80 ) a nd 7 1 ø .2 s yd (f rid er ich se n & b ro m ley 1 97 6) . f or le ge nd , s ee f ig . 2 18 (1988) for the bathonian, and summarised in callomon (2003). there is no a priori reason to believe that the faunal horizon-density, i.e. the reciprocal of the mean timeinterval between effectively instantaneous faunal horizons (callomon 1995, p. 136), is the same in the two scales. it is largely coincidental that the lowest horizon on the right in fig. 8 falls not far above its probable time equivalent in the left column, in the upper part of the humphriesianum standard zone. it does however suggest that the rate of morphological evolution of the boreal ammonites was not so very different from that of their cousins (at family-group level in the linnéan hierarchy of taxonomy) in the subtethyan domains farther south (callomon 1985a). but it does suggest that the records of the succession in the two domains are now known with comparable completeness. as knowledge of the greenland succession grew and there was no way of dating the pre-callovian part more closely, this part was comprehensively referred to at stage level simply as ‘boreal bathonian’ (callomon 1975, 1979, 1985a, b, 1993). although it is now certain that the lower part is in fact bajocian, the bajocian–bathonian boundary can still not be precisely located in the arctic. it cannot however lie far from the boundary between the pompeckji and arcticus standard zones (horizons 8/9), a boundary marked by an easily and widely recognisable morphological change in the ammonites. discoveries on the russian platform, near 733 1072 84 83 82 85 86 ørsted dal claudius clavus bjerge 5 km roll ier bjerge 72°00'n 22°50w23°00w23°10w23°20w23°30w kong oscar fjord 862 961 838 918 1138 1155 1093 841 1066 703 931 888 856 779 antarctic havn 71°50'n 71°55'n fig. 7. geological map of the area around antarctic havn showing locations of sections 82–86 of callomon (1993). map section from geological map sheets 71 ø.1 nord (perch-nielsen et al. 1983) and 72 ø.2 (escher 2001). for legend, see fig. 2. 19 saratov on the volga river (52°n, 500 km se of moscow; mitta & seltzer 2002), include arcticoceras of the basal ishmae standard zone (horizon 14) associated with parkinsonia ex gr. wuerttembergica characteristic of the top of the european lower bathonian zigzag standard zone (fig. 8). no great error is therefore likely to result for general purposes from assigning the borealis–pompeckji zones – the cranocephalites beds – to the upper bajocian and the arcticus–calyx standard zones to the bathonian. more recently, however, mitta (2009) argued that the green landicus zone in the timan–petchora region correlates with the central russian (volga) lowermost bathonian besnovi zone, which would then place the underlying arcticus zone in the uppermost bajocian, but this seems to need further confirmation. 1994–1996: ugleelv, central jameson land in compiling the faunal succession of fig. 8, the greatest difficulty lay in the lower part, in the indistinctus–pompeckji standard zones of the cranocephalites beds. the arcto cephalites and arcticoceras beds appeared to be widely repup pe r m id dl e lo w er up pe r m id dl e lo w er up pe r low er (p ar s.) ba th on ian c all ov ian zones zones & horizonszones lamberti athleta coronatum jason orbis zigzag parkinsoni garantiana niortense calloviense koenigi herveyi discus nordenskjoeldi apertum calyx variabile cranocephaloide ishmae greenlandicus arcticus pompeckji indistinctus borealis lamberti (alligatus) athleta coronatum anceps gracilis herveyi (bullatum) retrocostatum julii anguli costatum histricoides zigzag aurigerus bremeri morrisi subcontractus progracilis (as subboreal) (a) (submediterranean province not differentiated) banksi polygyralis blagdeni humphriesianum romani pinguis humphriesianum baculata dichotoma garantiana tetragona acris truellei bomfordi convergens macrescens yeovilensis tenuiplicatus progracilis subcontractus morrisi hodsoni hollandi discus keppleri terebratus kamptus gowerianus "curtilobus" galilaeii calloviense enodatum medea jason obductum grossouvrei phaeinum proniae spinosum henrici lamberti subzones & horizons subzones & horizons boreal province nw european / subboreal province submediterranean province b a c xx xix xviii xvii xvi xv xiv xiiib xiiia xii xi xb xa ix viii viib viia vi v iv iii ii i 14 15 16 13 12 11 10 9 1 2 3 4 5 6 7 8 j-37 j-36 j-35 j-34 j-33 j-32 j-31 j-30 j-29 j-24–26 j-28 j-23 j-22 j-21 j-20 j-19 j-18 j-17 j-14 j-13 j-11 j-9–10 j-8 j-3 j-2 j-1 ba jo cia n (p ar s.) fig. 8. ammonite zonal correlation chart with the primary european standard (column a, revised), and the secondary boreal standard (c) (callomon 1993, fig. 2, 2003, fig. 3, revised according to mangold 1990, 1997; mitta 2002, 2008, 2009; mönnig 2010, 2014). ammonite faunal horizons j1 – j37 recognised in the middle jurassic in jameson land up to 1993 are shown. they are given the prefix j to clearly distinguish them from the middle–upper jurassic succession of faunal horizons in nearby milne land (faunas 1–47 of callomon & birkelund 1980, 1982; birkelund & callomon 1985) which are accordingly labelled m1 – m47 (e.g. larsen et al. 2003). 20 resented in jameson land and elsewhere in the arctic, and their ammonite faunas well sampled in the collections from hurry inlet made by rosenkrantz in 1926–1927 and h. aldinger and others in later years. however, what was known of the cranocephalites beds was based largely on a single collection made by harris during an excursion in 1927 up the valley of ugleelv, west of upper hurry inlet, to its headwaters around katedralen (see map in rosenkrantz 1934, p. 86, fig. 34). this is also the type area of the oldest named species brought back in 1900, cranocephalites pompeckji (madsen 1904). spath’s description (1932) of harris’ material revealed a diverse spectrum of forms wholly unknown from hurry inlet but also represented in part in the collections made by rosenkrantz around mikael bjerg in north-central jameson land (fig. 4) where a caption to a photograph reads “in the foreground thousands of ammonites lie spread over the ground” (rosenkrantz 1929, p. 145). reconnaissance traverses by j.h. callomon in 1958 and during mapping by the university of copenhagen in 1970–1974 confirmed the richness of the faunas but could not resolve the stratigraphical uncertainties. each section seemed to have a specific, but different set of assemblages, the relative temporal relationships of which to each other remaining unclear. the faunal succession in the indistinctus and pompeckji standard zones was the best that could be done with the available evidence (callomon 1993; fig. 8, faunas j-2 – j-8). opportunities arose in 1994 and 1996 for callomon to revisit the area at ugleelv, in order to concentrate on the ammonite biostratigraphy of the cranocephalites beds. the results obtained exceeded all expectations. some 750 ammonites were collected bed-by-bed from 23 sections. the number of faunal horizons that could be distinguished grew from the nine recognised in 1993 (fig. 8) to thirtyfour today. it should be stressed however that the assemblages characterising these horizons differ widely in quantity 4 km 510 450 561 390 100 e1 e2 e3 d9 d10 d8 d7 d6 d5 d4 f d2 d1 c4 c3 c1 b2 b1b3 b4 a c2 d3 600 600 270 360 480 ræveelv ugleelv fa lk ee lv 430 480 610 620 450 380 500610 teebjerg 'taubjerg'statuebjerg katedralen n sortehat fig. 9. geological map of the area around ugleelv (see fig. 3), showing the locations of the 23 sections that have been recorded there up to 1996. they include sections 13–24 of callomon (1993), but they have been collectively newly labelled in six groups a–f as shown. for the correspondence between old and new numbers, see table 1; for legend, see fig. 2. note that ‘taubjerg’ is an informal name. 21 and quality. even when not sufficiently well-characterised to define a new nominal species, an assemblage may, however, suffice to affirm that it differs from others. the area around katedralen at ugleelv now provides the most detailed and complete biostratigraphical record of the ammonites of the cranocephalites beds anywhere and it is the purpose of this paper to describe this record. the ugleelv sections thus set the standard of reference, and records from other localities will be drawn in as secondary evidence and for comparison. further discoveries have also been made in subsequent years from scattered localities to the north of kong oscar fjord (fig. 1a; alsen & surlyk 2004; piasecki et al. 2004a, b) but these have not added to the list of faunal horizons. localities significant information on the cranocephalites beds in jameson land has come from three areas: (a) the coastal cliffs of neill klinter, along the western shores of hurry inlet, sections 1–8 of 1993 (fig. 2). (b) ugleelv, around katedralen and sortehat, sections 13–23 (fig. 3). (c) east-central jameson land, around centralbjerg, section 35, including hjørnefjeld of rosenkrantz (1929) and spath (1932), and trefjord bjerg, section 39 (fig. 4). by far the most extensive and detailed information has come from the ugleelv area, and this is taken to set the standard of reference. additional information has come from scattered localities to the north, at antarctic havn (figs 1b, 7; sections 82–86 of callomon 1993), traill ø on the northern side of kong oscar fjord (donovan 1953; alsen & surlyk 2004) and geographical society ø (fig. 1a; alsen 2015, appendix 1, this volume), but these will be only briefly alluded to. the locations of the 23 sections that have been studied in the ugleelv area are shown on the geological map in fig. 9. a new and independent section-numbering scheme is introduced for this key region, which prior to the 1994–1996 fieldwork included 12 sections (13–24 on figs 1b, 3; callomon 1993). table 1 illustrates the correspondence between the sections of callomon (1993) and the section series presented here for the key area of ugleelv (fig. 9). the sections lie in clusters, which may be conveniently labelled by a letter as prefix: a: ‘taubjerg’ (field name of callomon) at the eastern most outcrop. ‘taubjerg’ takes its name from the striking resemblance of the pelion formation outcrop seen in map view to the greek letter ‘τ’. b1–b4: statuebjerg (named by t.m. harris in rosenkrantz 1934, fig. 38). c1–c4: teebjerg (‘tee’ as in golf, named by t.m. harris in rosenkrantz 1934, fig. 38). d1–d10: katedralen (named by t.m. harris in rosen krantz 1934, fig. 38). e1–e3: ræveelv (named by t.m. harris in rosenkrantz 1934, fig. 38). f: sortehat (named by t.m. harris in rosenkrantz 1934, fig. 38). sections are subdivided into numbered beds with each section having separate bed numbering. the subdivision is based on field-observed surfaces. a bed numbered 1 in a given measured section is the lowest bed in that particular section, with the chosen base of section depending on the topography, outcrop exposure etc. thus bed 1 in one section does not necessarily, and is rather unlikely to, correlate with bed 1 of other sections. references to bed numbers in sections thus mainly serve to relate observations to a height or interval of those sections, typically the records of a taxon within the study area. lithostratigraphy the standard zone concept used by most jurassic ammonite workers is used throughout this paper also in the descriptions of lithostratigraphic units (see e.g. callomon 1985b, 1995, 2001; page 2003). it is important to stress that standard zones are chronozones and not biozones, even if their names are derived from the names of ammonite species. this significant difference, which is commonly overlooked, is indicated by using the non-italicised species name with an upper case first letter (e.g. borealis standard zone and not stratigraphy 22 the cranocephalites borealis zone or biozone). a standard zone can thus include unfossiliferous strata and its boundaries are isochronous surfaces in contrast to those of biozones. the middle jurassic ammonite-based zones of the present study are thus described as standard zones following the use of callomon (1985b, 1995, 2001), even if in the opinion of pa and fs they are biozones. jameson land the middle jurassic of jameson land is subdivided into the lithostratigraphical framework figured by surlyk (2003); formal lithostratigraphic revision of the jurassic of east greenland, as figured provisionally in surlyk (2003), is in progress. the middle jurassic succession is placed in the vardekløft group which was originally introduced as a formation by rosenkrantz (1929). it was revised, described and subdivided into the sortehat, pelion and fossilbjerget members by surlyk et al. (1973) who also gave a historical account. the sortehat member was excluded from the formation by surlyk (1991) and subsequently re-assigned at the status of formation to the newly-defined neill klinter group (dam & surlyk 1998). the upgraded vardekløft group includes the bristol elv, bastians dal, muslingebjerg, pelion, charcot bugt, fossilbjerget, payer dal, olympen, and jakobsstigen formations (surlyk 2003, fig. 5). in the present context, only those units occurring in jameson land are described, viz. the pelion, fossilbjerget and olympen formations. the group forms the top unit in most areas of its distribution so the maximum thickness is not well known, but is probably up to about 650 m thick. in jameson land, it comprises the sandstone-dominated pelion formation and the overlying mudstone-dominated fossilbjerget formation. the pelion formation overlies black mudstones of the aalenian – lower bajocian sortehat formation with a sharp boundary (dam & surlyk 1998; koppelhus & hansen 2003) and is overlain by black mudstones and massive sandstone injectites of the hareelv formation (surlyk et al. 2007). it was erected as the middle member of the vardekløft formation, which originally included the sortehat, pelion and fossilbjerget members (surlyk et al. 1973) and was preliminarily elevated to formation rank by surlyk (2003, fig. 5). key sections occur at pelion (type locality), gonio myakløft, zackenberg, katedralen, mikael bjerg, trefjord bjerg and olympen (figs 1–4). the thickness increases northwards from about 10 m in south-eastern jameson land to 310 m at the type section at pelion and 650 m at antarctic havn, although the latter section may include the lower part of the olympen formation. cross-bedded, planar-bedded and hummocky cross-stratified medium-grained, micaceous sandstones dominate, with subordinate coarse1994 field nos 1996 field nos goniomyakløft 1 fig. 16 'taubjerg' 96/4 a fig. 17 statuebjerg n 13 b1 fig. 18 96/1 b2 fig. 19 w 14 b3 sw 15 b4 a 16 94/18 c1 teebjerg b 17 94/18 c2 18 94/18 c4, 4' fig. 21 96/13 c3 fig. 20 katedralen e 19 d1 94/14 96/11 d2, 2' fig. 22 96/10 d3 96/2 d4 fig. 23 n 20 96/3 d5 fig. 24 96/6 d6 96/5 d7 94/15 d8 fig. 25 nw 21 d9 96/7 d10 fig. 26 94/16 e1 fig. 27 ræveelv a 22 94/13 e2 fig. 28a, bugleelv– b 23 94/17 96/4 e3 fig. 35a–c sortehat 24 96/12 f trefjord bjerg 39a 96/15 fig. 29 locality table 1. previous section labels in relation to this study callomon (1993, fig. 4) this study 23 grained and pebbly sandstones. the fossil content comprises ammonites, belemnites, bivalves, crinoids and brachiopods and deposition took place in marine shoreface and shallow shelf environments (for details, see engkilde & surlyk 2003). the formation spans the upper bajocian – middle callovian borealis–jason standard zones (faunal horizons 1–36 of callomon 1993), based on ammonites. the uppermost strata are poorly fossiliferous and the age relations to the overlying olympen formation are not always certain. in jameson land, the pelion formation is locally subdivided into the ugleelv and parnas members. the ugleelv member was named after the river in south-eastern jameson land (figs 1b, 3). the ugleelv member of the pelion formation overlies dark silty mudstones of the sortehat formation and is overlain by grey siltstones of the fos silbjerget formation. the upper boundary is a marked am monite-strewn drowning surface. the member is restricted to an area around inner ugleelv and ræveelv, south-eastern jameson land (fig. 3). it belongs to the upper bajocian borealis – basal pom peckji standard zones (faunal horizons 1–4 of callomon 1993), based on ammonites. the lower part of the member consists of sandstone (borealis standard zone), and the middle part of siltstone (in distinctus standard zone). the member is characterised by one (at katedralen) or two (at ræveelv) prominent, cliffforming, high-angle clinoform-bedded sandstones, which form the upper part of the member (see surlyk et al. 1973, fig. 19; heinberg & birkelund 1984, figs 13–14). the lower part of the member at these localities consists of structureless or cross-laminated, cross-bedded and hummocky cross-stratified, very fine-grained to medium-grained sandstones. it contains abundant ammonites and belemnites, bivalves, crinoids, brachiopods, rare gastropods, trace fossils and wood fragments. fossils are commonly concentrated in distinct layers. deposition took place in shallow marine shelf and shoreface environments. the upper clinoform-bedded part of the member was deposited in a tidally-influenced shoreface transition zone. the overlying parnas member was first recognised as a sandstone wedge in the top part of the pelion formation by heinberg & birkelund (1984). the member is named after the mountain parnas in north-central jameson land and the type section is on the north-east slope of the mountain olympen, also in northern central jameson land (fig. 1b). the member is about 30 m thick and consists of very fine-grained to coarse-grained sandstones forming metrescale coarsening-upward cycles. it contains abundant ammonites and belemnites, rare bivalves, trace fossils and wood fragments, and deposition took place in a shallow marine shelf to shoreface. it interdigitates with the upper part of the fossilbjerget formation, hence temporarily interrupting the general northwards backstepping nature of the pelion–fossilbjerget boundary. it overlies very finegrained silty sandstones of the fossilbjerget formation with a gradational transition. the upper boundary is a major drowning surface, which is sharply overlain by very finegrained silty sandstone of the fossilbjerget formation. the parnas member is exposed in the mountains of olympen, parnas, pelion and on the north slopes of fossilbjerget in northern central jameson land. it wedges out towards the south and is not present at mikael bjerg (figs 1, 4). it belongs to the lower callovian apertum–nordenskjoeldi standard zones (faunal horizons 24–30 of callomon 1993), based on ammonites. the fossilbjerget formation was erected as a member of the vardekløft formation by surlyk et al. (1973) and was preliminarily elevated to formation rank by surlyk (2003, fig. 5). the type section is at the mountain of fossilbjerget (surlyk et al. 1973, fig. 23; fig. 1b) and key sections occur at goniomyakløft, zackenberg, katedralen, mikael bjerg, and pelion (figs 1–4). the formation is 80–120 m thick and consists of silty, highly micaceous mudstone, with subordinate beds of fine-grained sandstone. ammonites occur in profusion, and bivalves, belemnites, dinoflagellate cysts, tree trunks and trace fossils are common. deposition took place in an offshore marine environment. the formation overlies sandstones of the pelion formation with a highly diachronous boundary, younging towards the north and is overlain in central jameson land by massive sandstones of the olympen formation (athene member) and in southern jameson land by a thin unit of black mudstones of the olympen formation (hades member). it occurs throughout jameson land but appears to have been removed by modern erosion in the northernmost part. as noted above, the sandy parnas member of the pelion formation forms a southward-tapering wedge in the top part of the formation in northern central jameson land (fig. 1b).the fossilbjerget formation belongs to the upper bajocian – lower upper callovian, pompeckji–athleta standard zones (faunal horizons 3–37 of callomon 1993), based on ammonites. the goniomyakløft member was recognised as a sharpbased unit of soft light-grey weathering shales, about 20 m thick, forming the top of the vardekløft formation in southern jameson land (surlyk et al. 1973). the member corresponds roughly to the ‘wood beds’ of callomon (1993). the type section is at goniomyakløft, hurry inlet (section 6, fig. 14 in surlyk et al. 1973) and good sections occur at zackenberg and katedralen (fig. 2). the member consists of light-grey mudstones with large, greenish, irregular, extremely hard concretions commonly with a nucleus formed by silicified logs. rare ammonites are found 24 in the concretions. deposition took place in an offshore marine environment. the member is about 20 m thick and overlies silty, dark-grey mudstones with thin finegrained sandstone intercalations of the lower fossilbjerget formation with a sharp contact corresponding to a major hiatus. it is overlain by black mudstones or massive sandstones of the olympen or hareelv formations. it is restricted to south and southern central jameson land. the member belongs to the top lower callovian – lower upper cal lovian, calloviense–athleta standard zones, based on ammonites. the pelion and fossilbjerget formations exhibit major shifts in facies, in some cases associated with non-sequences of considerable magnitude, reflecting important regional changes in relative sea level and sedimentary regime. characteristically, the succession consists of soft, barely cemented recessive beds of uniform siliciclastic lithology from decimetres to metres thick, punctuated by thin resistant concretionary layers comprising small or mediumsized, hard, calcareous or sideritic concretions and large sandstone ‘doggers’ (term much used by callomon to describe ellipsoidally weathering sandstone bodies that are more strongly cemented than the otherwise identical adjacent sandstone). these thin, preferentially cemented beds serve as invaluable markers for correlation. they extend over distances of just hundreds of metres to tens of kilometres. they are commonly marked by concentrations of glauconite and by ferruginous or phosphatic cementation, giving them reddish or brown colours when weathered. they are also the main sources of well-preserved body fossils, particularly ammonites, which can occur in spectacular assemblages of complete, beautifully preserved adults strongly sexually polarised in favour of (female) macroconchs (see, for example, figs 34, 38, 39, 42, 43; cf. callomon 1963). among representatives of the nekton, belemnites can also be abundant. the benthos is represented by sparse faunas of bivalves, gastropods, terebratulid brachiopods and occasional small solitary corals and crinoids. statuebjerg teebjerg katedralen vardekløft group neill klinter group hall bredning group ugleelv c2 c1 hall bredning group b1 fig. 10. aerial view of the southern valley side of ugleelv, looking almost due west. the highest peak in the far distance on the right is katedralen. coming nearer, the two long ridges falling into the valley are teebjerg and statuebjerg respectively, rising in the far distance on the left into the regional plateau formed by the resistant hareelv formation (hall bredning group). the scarp-edged plateau in the near distance is formed by the top of the resistant sandstones of the neill klinter group, rising via the recessive sortehat and lower pelion formations (vardekløft group) to another small plateau, ‘taubjerg’ (field name of callomon), at the foot of the final rise to the hareelv plateau. locations of sections b1, c1 and c2 indicated. 25 there is a diverse ichnofauna in many beds (heinberg & birkelund 1984). plant remains are abundant, mostly in comminuted form, and large drifted logs of wood are commonly found as cores of siliceous concretions in the higher parts of the fossilbjerget formation, notably in the goniomyakløft member. the olympen formation, which is missing in southern jameson land, is first recognised at langryggen and thickens northwards to 180 m at olympen (sections 27 and 67 of callomon 1993, respectively; fig. 1b). the formation belongs to the lower to middle oxfordian. the hareelv formation comprises black mudstones, and density-flow sandstones which have been remobilised and injected into the surrounding mudstones (surlyk & noenygaard 2001, 2003; surlyk et al. 2007). it forms the top plateau of much of southern jameson land (figs 10, 11). the formation belongs to the upper oxfordian – kim meridgian. ugleelv – northern hurry inlet the pelion formation occurs along the whole length of jurassic outcrops in central east greenland, from hurry inlet (71°n) to store koldewey (76°n). its age-range varies considerably from place to place. it is subdivided into a succession of members (surlyk 2003, p. 666, fig. 5), but of these only the lowest, the ugleelv member, is of interest here. this encompasses all the beds of the pelion formation found in the area around ugleelv and then southwards along neill klinter. the ugleelv member can be further sortehat ugleelv teebjerg katedralen vardekløft group vardekløft group ræveelv hall bredning group (hareelv formation) d10 c1 c2 c4 d9 e2 e3 d2 d3c3 neill klinter group d4 d5 d6 d7 fig. 11. aerial view of ugleelv and katedralen seen from east-north-east, looking towards the south-west on the horizon to the plateau basalt cliffs on the south side of scoresby sund (left) and the caledonian crystalline basement of milne land (right). in the foreground, a thin capping of lower pelion formation on sortehat formation (steeper slope), resting in turn on the platform top of the neill klinter group (ostreaelv formation (toarcian)). the highest peak in the line of sight is katedralen, its flat top formed by the hareelv formation. the ridge leading from its peak downwards to the left passes through sections d2–d4. the ridge leading downwards to the right, in front of the canyon, is section d5, the most prolific source of cranocephalites pompeckji. this major, deeply incised canyon is flanked by sections d8 (not visible) and d9. the prominent step half-way up marks the clinoforms at the top of the pelion formation. the ridge from the plateau at far right, leading into the headwaters of ugleelv, is section d10. the ridge on the northern (near) side of ugleelv, facing the katedralen canyon and topped by the clinoforms, hosts sections e2 and e3. the rounded hill in the far foreground, in direct line to katedralen, is sortehat itself, section f (not visible). 26 400 300 200 100 ostreaelv fm 0 m ammonite legend for figs 12, 16–29, 35 fossils trace fossils structures, lithologies belemnite belemnoteuthid cephalopod gastropod pectinid bivalve bivalves, others terebratulid brachiopod diplocraterion habichi plant massive sandstone log carbonate concretion metres above sea levelm a.sl. clinoform bedded sandstone sandstone with trace fossils concretionary sandstone (dogger) siltstone sandstone black mudstone phosphatic concretion phoebichnus trochoides lithostratigraphy u gle el v m em be r u g1 u g2 sh -1 sh -2 u g3 fb -2 fb -1 u g4 fb -3 fb -4 fo ss ilb je rg et f or m at io n h ar ee lv fm h all b re dn in g g r. o f pe lio n fo rm at io n so rt eh at f or m at io n n ei ll kl in te r g ro up va rd ek lø ft g ro up fig. 12. a synoptic diagram of the development of the vardekløft group at katedralen and its surroundings. the succession is based on weathering profiles drawn by jhc in the field (see figs 16–29, 35). they mainly illustrate differences in consolidation and cementation of the beds and have less emphasis on the sedimentary structures. the legend also applies to figs 16–29 and 35. informal lithostratigraphic units within the sortehat formation (sh-1, sh-2), the pelion formation (ug-1 – ug-4) and the fossilbjerget formation (fb-1 – fb-4) are indicated. note the diachronous nature of the pelion–fossilbjerget boundary. of: olympen formation. 27 subdivided in the region between ugleelv and hurry inlet into roughly four parts that can be widely discerned there. the lithological differences are small, however, and do not merit formal lithostratigraphical status; they are informally labelled ug-1 to ug-4 from below (fig. 12). ugleelv unit 1 (ug-1). light-coloured quartzose sandstones, heavily burrowed, and rich in plant remains, without marine macrofossils other than rare belemnites throughout. thicknesses 20–70 m. strontium isotope dating of the belemnites points to an age around the early/late bajocian boundary (m. engkilde, personal communication 1998). ug-2. light-coloured sandstones with subordinate silty shales and thin, lenticular interbeds of ironstone, probably of glauconitic origin, commonly associated with small phosphatic concretions and concentrations of fossils, including ammonites, belemnites (sometimes as ‘battlefields’, fig. 13), bivalves and gastropods. the unit is 25–50 m thick 10 cm fig. 13. a belemnite ‘battlefield’ at a thin interbed with phosphatic concretions near the top of ug-2 (pelion formation), section e3 (fig. 9, see also fig. 35b). hareelv formation pelion formation pelion formation sortehat formation fossilbjerget formation hareelv formation fossilbjerget formationd7 d4 d5 d8 fig. 14. the northern slopes of katedralen seen from the north-west, viewed from section d9 (fig. 9). in sections d7 and d8, the ferruginous cranocephalites beds (fossilbjerget formation, fb-1) are only 2–3 m thick and rest immediately on the clinoforms (pelion formation, ug-4), which are up to 25 m thick. the overlying brown-coloured shales of fb-2 are essentially the arctocephalites and perhaps lower arcticoceras beds. the light and darkercoloured shales that follow in the upper fossilbjerget formation are fb-3 and fb-4 respectively (fig. 12). 28 and terminates around katedralen fairly sharply at one of these horizons of ironstone and phosphatic concretions, the latter commonly enclosing a large plagiostoma bivalve, making an excellent marker for local correlation in the ugleelv area. ug-2 belongs to the borealis standard zone (all three horizons, bo-1 – bo-3, see below). ug-3. siltstones to fine-grained muddy sandstones with occasional thin lenticular concretionary, slightly ferruginous layers containing scattered light-brown weathering doggers, which are the only source of recoverable body fossils. the unit coarsens upwards into sandstone. ug-3 belongs to the indistinctus – lower pompeckji standard zones. it is typically about 20 m thick around katedralen, but thins eastwards to zero at ‘taubjerg’ (fig. 9, section a) and grades upwards at katedralen into the overlying unit ug-4. ug-4. the large-scale, high-angle clinoform sandstone beds at katedralen (upper part of the ugleelv member of surlyk (2003, fig. 5) which attains thicknesses of up to 30 m (fig. 14). it is variably divisible into subsets, in some cases separated by softer, muddy interbeds that may yield ammonites. it thins and wedges out eastwards, grading laterally into the upper part of ug-3. at its thickest, the top surface is sharp and heavily pierced by vertical burrows of diplocraterion habichi and possibly also monocraterion. the unit belongs to a single ammonite biohorizon in the pompeckji standard zone, po-4. the regional variations in the thickness of the pelion formation are illustrated in fig. 15. the pelion formation is sharply overlain by the fos silbjerget formation which consists predominantly of mudstones punctuated by layers of concretions or concretionary induration. the gross lithologies are constant over large distances with relatively small changes in thickness. a subdivision into four parts can again be recognised in the area around ugleelv and katedralen. the fossilbjerget formation is subdivided into units fb-1 to fb-4. fossilbjerget unit 1 (fb-1). a highly condensed succession of thin but sharply bounded ironstones, weathering 100 m 50 0 d10 sections d8 e3 d7d9 e1 d6 e2 fb-4 fb-4 fb-3 fb-2 fb-2 ug-3 ug-1, ug-2ug-1, ug-2 fb-4 fb-3 fb-2 fb-1 ug-4 ug-3 ug-3 ug-4 ug-4 ug-3 ug-3 ug-3 fb-1 fb-2 fb-3 fb-4 -27 km024681012 d5 d4d3 f d1 c3-4 hareelv formation fossilbjerget formation pelion formation sortehat formation c1c2 b3-4 b2b1 a 1d2 nw se s fig. 15. regional variations in the development of the pelion and fossilbjerget formations around ugleelv. selected sections are projected onto a gently curving arc as baseline from section a in the south-east to section d10 in the north-west (fig. 9). numbers in columns refer to the informal subdivision of the pelion formation into ug-1 – ug-4 and the fossilbjerget formation into fb-1 – fb-4. the lower two pelion units, ug-1 and ug-2, are shown undivided. ug-4: clinoforms. the diagram illustrates that the pelion formation becomes thicker and stratigraphically more complete towards the north-west. 29 red-brown, separated by soft muddy interbeds. westwards from hurry inlet, they onlap against and then overlap the clinoforms of ug-4, so that on top of it the ironstones are in contact. the ironstones are locally highly fossiliferous, especially with ammonites, whose uncrushed body chambers are commonly also somewhat phosphatised. the ammonites have allowed us to establish detailed age relationships, showing that the unit represents a very long period of time with up to 15 ammonite horizons in 5 m of sediment. individual beds can wedge out over small distances and the succession in any one section can be full of non-sequences. the unit belongs to the pompeckji standard zone, horizons po-8 to po-23 at katedralen, and the lower arcticus standard zone, a. arcticus horizon j9, from north of ugleelv (section e3) as far as mikael bjerg and centralbjerg, 35 km to the north (fig. 1b). farther northwards, it interdigitates with the pelion formation. the unit as a whole wedges out eastwards and is completely missing in a major non-sequence at section a (‘taubjerg’; fig. 9), and then southwards along the whole of neill klinter. this non-sequence may in part be erosional for it is marked by a sharp contact with the overlying beds. the unit is up to 5 m thick, before grading into the pelion formation. fb-2. dark silty mudstones with well-separated thin layers of ferruginous induration or flat concretions in the lower part weathering reddish-brown, changing to lighter mudstones with layers of fine-grained sandstones or fissile concretions in the upper part. these are respectively the arctocephalites and arcticoceras beds (spath 1932). the slopes of katedralen mark the type locality of the greenlandicus standard zone. at neill klinter, the beds are highly fossiliferous and have yielded large collections of the ammonites giving name to the beds. at katedralen, they are only sparsely fossiliferous. some of the beds in the greenlandicus standard zone are marked by spectacular examples of the large, wheel-like burrows phoebichnus trochoides. the unit belongs to the greenlandicus to ishmae standard zones. it is 50 m thick at hurry inlet, thinning westwards to 30 m at katedralen and merging with the pelion formation in the north. fb-3. silty pyritic mudstones, with only a few indurated layers and scattered, hard, sideritic spherical concretions, the kepplerites–cadoceras beds (spath 1932). the unit terminates with a widespread bed of concretionary calcareous siltstone or fine-grained sandstone, the calyx bed at hurry inlet, an excellent marker still recognisable in ugleelv. two key sections at fossilbjerget, nos 42 and 43 in fig. 4, were described by callomon (2004). the unit belongs to the upper boreal bathonian cranocephaloide–calyx standard zones from hurry inlet to katedralen. it is around 30 m thick, changing little laterally. fb-4. this unit has previously been termed the ‘wood beds’ (callomon 1993) and is now referred to the goniomyakløft member (surlyk 2003, fig. 5). it comprises fine-grained monotonous mudstones, with only sparse concretions and rare macrofossils but abundant silicified wooden logs. the unit belongs to the lower callovian calloviense standard zone to the upper callovian. the thickness of c. 50 m is very persistent laterally. it terminates in southern jameson land in a non-sequence that cuts out the equivalents of the lower oxfordian part of the olympen formation. the intricate stratigraphic relationships of the ammonite faunal horizons in the ugleelv region, and their position within the lithostratigraphic framework described above, were determined by detailed collecting linked to stratigraphic logging of 23 sections (fig. 9). key sections, mainly from the ugleelv area, are illustrated here in figs 16–29. 30 450 2 3 4 5 6 7 8 9 10 12 14 20 22 24 26 33 11a 500 550 600 650 11b 11c 11d cadoceras calyx, kepplerites vardekloeftensis kepplerites rosenkrantzi kepplerites cf. tychonis cadoceras kepplerites stephanoides arcticoceras crassiplicatum arcticoceras ishmae arctocephalites delicatus cranocephalites borealis γ goniomyakløftlithostratigraphy ammonite m a . s l. be d no . g on io m ya kl øf t m b. fo ss ilb je rg et f or m at io n h ar ee lv fo rm at io n so rt eh at f m o s f m pe lio n fo rm at io n fig. 16. section from goniomyakløft (section 1 of callomon 1993; for location, see fig. 2); for legend, see fig. 12. os fm: ostreaelv formation. 31 450 3 12 4 6 10 12 16 18 20 22 24 26 28 30 32 34 36500 400 'taubjerg': section alithostratigraphy ammonite sample no. m a . s l. be d no . u gle el v m em be r u g1 u g2 fb -2 fb -3 fb -4 pe lio n fo rm at io n so rt eh at fo rm at io n fo ss ilb je rg et f or m at io n [ug-3, 4, fb-1] cadoceras sp. cf./aff. nordenskjoeldi jhc 6583–6587 kepplerites traillensis β, cadoceras apertum β kepplerites vardekloeftensis, cadoceras calyx (not collected) arctocephalites greenlandicus β arctocephalites (not collected) arcticoceras ishmae arcticoceras crassiplicatum arcticoceras crassiplicatum kepplerites peramplus (not collected) jhc 6574 jhc 6575–6577 arcticoceras ishmae β arcticoceras cf. harlandi jhc 6564–6573 jhc 6555–6563 (not collected) jhc 6546–6554 jhc 6578–6582 fig. 17. section a at ‘taubjerg’. for location, see fig. 9; for legend, see fig. 12. note that the ug-3, ug-4 and fb-1 units are absent at the pelion–fossilbjerget boundary. 32 350 2 34 5 7 8 9 10 11 13 14 15 16 18 19 20 21 23 22 24 25 26 27 28 29 30 31 32 3435 33 36 37 38 39 40 41 42 43 44 45 46 48 47 49 50 52 53 54 55 51 6 costacadoceras jhc 1235 kepplerites, cadoceras? arcticoceras harlandi (j-14) jhc 1232–1234 arctocephalites jhc 1231 arctocephalites jhc 1231 cranocephalites carlsbergensis γ (po-12) jhc 1249–1279 cranocephalites indistinctus β (in-6) jhc 1312–1319 cranocephalites indistinctus α (in-3) jhc 1248, 1311 arctocephalites greenlandicus (j-12–j-13) jhc 1226–1230 kepplerites cf. rosenkrantzi (j-21) kepplerites [stephanoides], jhc 1237–1240 jhc 1246–1247 arcticoceras crassiplicatum (j-17–j-18) jhc 1241–1245 cranocephalites borealis β? (bo-2?) cranocephalites borealis β (bo-2) 400 12 17 450 500 550 600 statuebjerg n: section b1lithostratigraphy ammonite sample no. m a . s l. be d no . fb -4 fb -2 u g3 u g2 u g1 fb-1 fb -3 fo ss ilb je rg et f or m at io n pe lio n fo rm at io n so rt eh at f m h ar ee lv fo rm at io n g on io m ya kl øf t m b u gle el v m em be r fig. 18. section b1 at statuebjerg n, with samples from before 1994 indicated. for location, see fig. 9; for legend, see fig. 12. note, in comparison to fig. 17, that only the ug-4 unit is lacking at the pelion–fossilbjerget boundary, but over 10 faunal horizons are absent at this surface (see also fig. 19). 33 450 460 470 440 16 15 14 10 9 8 7 6 12a 11b 12b 13b 13a 11a cranocephalites carlsbergensis γ (po-12) cranocephalites borealis β (bo-2) cranocephalites borealis β (bo-2) cranocephalites borealis γ (bo-3) cranocephalites borealis cranocephalites indistinctus α (in-3) cranocephalites indistinctus cranocephalites indistinctus γ (in-8) cranocephalites indistinctus (in-7?) cranocephalites ex. gr. indistinctus (u-4) (in-7) cranocephalites indistinctus α' (in-4) jhc 6048–6070 (not collected) (not collected) jhc 6001–6006 jhc 6007–6016 (not collected) jhc 6030–6047 jhc 6025–6029 jhc 6023–6024 jhc 6017–6022 statuebjerg: section b2lithostratigraphy ammonite sample no. m a . s l. be d no . u gle el v m em be r fb -2 u g3 u g2 fb-1 fo ss ilb je rg et f or m at io n pe lio n fo rm at io n fig. 19. section b2 at statuebjerg; same interval as in nearby section b1 (fig. 18) but figured in greater detail and showing samples from the j.h. callomon 1996 collection. for location, see fig. 9; for legend, see fig. 12. 34 0 7 6 5 4 3b 2 3a 1d 1b 3c 10 20 30 cranocephalites intermissus α (po-3) cranocephalites sp. cranocephalites carolae α (po-1) jhc 6507–6530 (not collected) jhc 6493–6506 cranocephalites sp. (u-10) (po-23) teebjerg–falkeelv: section c3lithostratigraphy ammonite sample no. m et re s be d no . u gle el v m em be r fb -2 fb -1 u g3 u g2 fo ss ilb je rg et f or m at io n pe lio n fo rm at io n jhc 6531 fig. 20. section c3 at teebjerg–falkeelv. for location, see fig. 9; for legend, see fig. 12. 35 550 500 450 400 600 kepplerites tychonis, cadoceras variabile (j-21–j-23?) arctocephalites greenlandicus (j-12) jhc 5254–5268 cranocephalites borealis jhc 5202–5203 cranocephalites borealis γ (bo-3) jhc 5186–5201 cranocephalites sp. (u-10) (po-23) jhc 5251–5252 cranocephalites sp. cranocephalites sp. aff. gracilis (u-9) (po-16) jhc 5235–5250 cranocephalites carolae α (po-1) cranocephalites sp. jhc 5204–5234 (not collected) (not collected) cadoceras calyx cadoceras sp., kepplerites tenuifasciculatus (j-27) cadoceras cf. variabile, kepplerites sp. cadoceras cf. variabile? falkeelv–teebjerg: section c4+c4'lithostratigraphy ammonite sample no. m a . s l. be d no . u gle el v m em be r g on io m ya kl øf t m em be r fb -4 fb -3 fb -2 fb -1 u g2 u g3 u g1 h ar ee lv fm fo ss ilb je rg et f or m at io n pe lio n fo rm at io n 35 33 31 29 27 26 25 24 23 21 19 17 15 13 12b 12c 12d 12a 11 10 9 8 7 6 5 4 3 2 sortehat fm fig. 21. section c4, c4´ at falkeelv–teebjerg. for location, see fig. 9; for legend, see fig. 12. 36 420 410 400 390 4 3 2 1 5 6 7 8 9 10 11 13 14 15 16 17 18 19 20 12a 12b 380 375 dyke plagiostoma bed arctocephalites cf. arcticus (j-9) (not collected) cranocephalites sp. (u-10) (po-23) cranocephalites sp. ind. jhc 5431–5438, jhc 6481–6486 cranocephalites sp. (u-2) (in-2) jhc 5313–5322, jhc 6447–6457 cranocephalites sp. (u-1) (in-1) jhc 5306–5312 cranocephalites borealis γ (bo-3) jhc 5311 cranocephalites borealis γ (bo-3) jhc 5304–5305 cranocephalites carlsbergensis α (po-10) jhc 5391–5411 cranocephalites furcatus (po-9) jhc 5372–5390 cranocephalites carolae β (po-2) jhc 5324–5342, jhc 6458–6466 cranocephalites gracilis (po-15) jhc 5412–5430, jhc 6467–6480 cranocephalites indistinctus α (in-3) jhc 5323, jhc 5343–5371 katedralen e: section d2+d2'lithostratigraphy ammonite sample no. m a . s l. be d no . fo ss ilb je rg et f or m at io n u g3 u g4 fb -1 fb -2 u g2 u gle el v m em be r fig. 22. section d2, d2´ at katedralen e. for location, see fig. 9; for legend, see fig. 12. 37 0 12 da e d b a 10 9 8 6 5 4 3 2 1 150 100 50 cranocephalites aff. tvaerdalensis (u-8) (po-14) cranocephalites tvaerdalensis (po-13) cranocephalites indistinctus cranocephalites carlsbergensis α (po-10) cranocephalites gracilis (po-15) jhc 6132–6177 jhc 6114–6131 jhc 6085–6113 jhc 6072–6084 arctocephalites sp. (not collected) (not collected) katedralen n: section d4lithostratigraphy ammonite sample no. m et re s be d no . u gle el v m em be r u g4 u g3 u g2 u g1 fo ss ilb jer ge t f m pe lio n fo rm at io n so rt eh at f m fb-1 fig. 23. section d4 at katedralen n. for location, see fig. 9; for legend, see fig. 12. 380 370 360 350 z 6 a b c b a 5 4 3 2 1 390 cranocephalites gracilis (po-15) cranocephalites tvaerdalensis (po-13) cranocephalites pompeckji (po-8) cranocephalites intermissus β (po-4) cranocephalites sp. (u-2) (in-2) cranocephalites sp. (u-1) (in-1) cranocephalites sp. (u-1) (in-1) cranocephalites borealis γ (bo-3) jhc 6185–6191 jhc 6183, 6184 jhc 6179–6182 jhc 6178 cranocephalites maculatus (po-19) cranocephalites carlsbergensis (~po-10–po-12) (not collected) jhc 6193–6208 (not collected) cranocephalites episcopalis α (po-21) (not collected) jhc 6209–6234 jhc 6235 jhc 6192 katedralen nw: section d5lithostratigraphy ammonite sample no. m a . s l. be d no . u gle el v m em be r fb -1 fb -2 u g3 u g2 u g4 fo ss ilb je rg et f or m at io n pe lio n fo rm at io n fig. 24. section d5 at katedralen nw. for location, see fig. 9; for legend, see fig. 12. 39 310 300 290 280 270 18 a d b 17 16 15 14 13 12 11 335 (26 m) cranocephalites gracilis (po-15) cranocephalites borealis β (bo-2) plagiostoma (bivalve) jhc 5101–5107 cranocephalites episcopalis β (po-22) jhc 5158–5159 cranocephalites episcopalis α (po-21) jhc 5154–5157 cranocephalites ornatus (po-20) jhc 5146–5153 cranocephalites transitorius α (po-17) jhc 5138–5145 jhc 5128–5137 cranocephalites sp. (po-5) jhc 5125–5127 cranocephalites intermissus β (po-4) jhc 5113–5124 cranocephalites indistinctus γ (in-8) jhc 5108–5112 katedralen n canyon: section d8lithostratigraphy ammonite sample no. m a . s l. be d no . u gle el v m em be r u g4 u g3 u g2 pe lio n fo rm at io n fo ss ilb je rg et f m fig. 25. section d8 at katedralen n. for location, see fig. 9; for legend, see fig. 12. 40 500 450 400 350 300 30 29 28 27 21 14 12 11 10 9 8 7 6 5 4 3 2 1 22 20 15 13 24 26 fb-1 katedralen w: section d10lithostratigraphy ammonite sample no. m a . s l. be d no . fo ss ilb je rg et f or m at io n pe lio n fo rm at io n g on io m ya kl øf t m em be r u gle el v m em be r h ar ee lv fo rm at io n sigaloceras calloviense, pseudocadoceras (j-35) jhc 6421–6423 kepplerites vardekloeftensis, cadoceras calyx (j-23) jhc 6418–6420 arctocephalites greenlandicus (j-12) jhc 6491–6502 cranocephalites episcopalis α (po-21) jhc 6404–6405 cranocephalites spp. (not collected) cranocephalites gracilis (po-15) jhc 6410–6417 fb -4 fb -3 fb -2 u g4 u g3 u g2 u g1 cranocephalites indistinctus α (in-3) jhc 6403 cranocephalites borealis β (bo-2) jhc 6390–6402 cranocephalites borealis β (bo-2) jhc 6388–6389 cranocephalites borealis β (bo-2) jhc 6374–6387 cranocephalites borealis α (bo-1) jhc 6369–6373 fig. 26. section d10 at katedralen w (upper ugleelv). for location, see fig. 9; for legend, see fig. 12. 41 430 12 11 10 9 8 7 6 5 4 3 2 1 a b c a a b b d 420 410 400 390 380 cranocephalites intermissus β (po-4) jhc 5172–5185 cranocephalites ex. gr. indistinctus (u-3) (in-5) jhc 5166–5171 plagiostoma (bivalve) cranocephalites borealis β (bo-2) jhc 5160–5165 ugleelv n: section e1lithostratigraphy ammonite sample no. m a . s l. be d no . u gle el v m em be r u g4 u g3 u g2 pe lio n fo rm at io n 35 fig. 27. section e1 at ugleelv n. for location, see fig. 9; for legend, see fig. 12. 42 350 300 250 cranocephalites borealis α (bo-1) cranocephalites borealis α ? cranocephalites borealis β (bo-2) cranocephalites borealis β (bo-2) cranocephalites borealis β (bo-2) cranocephalites borealis β (bo-2) cranocephalites borealis β (bo-2) jhc 1373 jhc 1374–1376, jhc 5491 jhc 1377–1379 jhc 1380 jhc 1381 jhc 1382–1387 cranocephalites sp. cranocephalites borealis cranocephalites indistinctus β (in-6) cranocephalites intermissus β (po-4) jhc 1360–1372, jhc 5283–5303 jhc 1388 jhc 1389a, b, jhc 5492–5493 jhc 5494 ræveelv: section e2 a fig. 28b lithostratigraphy ammonite sample no. m a . s l. be d no . u gle el v m em be r u g3 u g2 u g1 u g4 pe lio n fo rm at io n so rt eh at f m 38 37 36 3534 30 33 32 31 29 28 27 26 13 12 11 10 9 8 7 6 222120 19 18 17 16 15 14 24 23 25 fig. 28. section e2 at ræveelv. for location, see fig. 9; for legend, see fig. 12. a: general section. b (facing page): detailed section of the upper part shown in a. 43 360 370 350 340 380 38 37 36 35 34 33 32 31 30 29 28 27 26 25 24 23 cranocephalites indistinctus β (in-6) cranocephalites intermissus β (po-4) jhc 1389a, b jhc 5492–5493 jhc 5494 jhc 1388 cranocephalites borealis cranocephalites sp. (not collected) plagiostoma (bivalve) cranocephalites borealis β (bo-2) jhc 1382–1387 cranocephalites borealis β (bo-2) jhc 1381 cranocephalites borealis β (bo-2) jhc 1380 cranocephalites borealis β (bo-2) jhc 1377–1379 ræveelv: section e2 blithostratigraphy ammonite sample no. m a . s l. be d no . u g3 u g2 u g4 pe lio n fo rm at io n u gle el v m em be r 3 4 5 6 89 10 13 12 11 14 15 16 17 18 19 20 21 22 23 7 2 1 800 900 950 1000 cranocephalites borealis α (bo-1) cranocephalites carlsbergensis s.s., β (po-11) jhc 1724–1768 jhc 6588–6590 trefjord bjerg: section 39a, blithostratigraphy ammonite sample no. m a . s l. be d no . pe lio n fo rm at io n so rt eh at f or m at io n o s f m fig. 29. section 39a, b attrefjord bjerg. for location, see fig. 4; for legend, see fig. 12. os fm: ostreaelv formation. the section profile is based on section 39b (fig. 40); specimens jhc 6588–6590 from section 39a are located by correlation. note that a significant hiatus is indicated between beds 10 and 11 where 20 faunal horizons are absent (fig. 30). biostratigraphy the ammonite biostratigraphy of jameson land, and indeed of the whole of east greenland, presents some special features. first, the fossils can, with some exceptions, be collected only from the hard, preferentially cemented beds. in the sandstones of the pelion formation, these are almost always well-bounded by clear partings, so that bed-by-bed collecting can be sharply defined. the casts of the shells are well preserved. in the softer mudstones of the upper pelion and fossilbjerget formations, the collectable ammonites are confined to the concretions or otherwise diagenetically indurated horizons that commonly mark the tops of sedimentary cycles. in some cases, the fossils form the nucleus of concretions, particularly in the case of ammonites. dia genesis was early, for the fossils are generally not distorted by compaction, and one of the striking features of the greenland collections is their fine state of uncrushed preservation. phragmocones of ammonites may still be void, or filled with calcite or, as witness perhaps to mild palaeogene metasomatism, single crystals of barytes. the shells are neither broken nor encrusted with epizoans. whether fossils also occur in the soft mudstones that separate the levels of concretions is not known, for the beds are either superficially weathered or impenetrable through permafrost. assemblages collected from such narrowly-defined stratigraphical horizons approach closely the ideal of what have been termed ‘faunal biohorizons’ (e.g. callomon 1964, 1985a, b, 1995), viz. assemblages within which no further biostratigraphical differentiation can be resolved and which must therefore be regarded as effectively instantaneously isochronous on the time scale of discernible ammonite evolution. a succession of such assemblages of a group, such as an ammonite genus, therefore gives a time-serial snapshot record of the group and, if its morphology changes with time, of its evolution. these biostratigraphical units recording unknown but brief periods of formation separated by equally unknown intervals of non-recorded time form the basic input of a bottom-upward synthesis of the palaeobiology of an evolving group (more detailed discussions in callomon 1985a, 1995). the second point to be noted is that the faunal horizons are highly impersistent. the hard beds come and go, both vertically and horizontally, over distances of as little as a kilometre. even when beds can be followed over greater distances, their quantitative fossil-content may change drastically. the outcome is that in constructing a correct timeordered sequence of faunal horizons, recourse has to be made to correlations between sections sufficiently close together geographically for their biohorizons to overlap laterally or in succession. the correlations are then usually lithostratigraphical, relative to selected markers. even so, uncertainties in relative positions of faunal horizons can remain, for in some cases a well-defined horizon has been found in isolation at only a single locality. the number of faunal horizons recognised in the cranocephalites beds of jameson land prior to 1994 was eight. they were based on material from scattered localities, with little direct evidence of the stratigraphical interrelationships. at that time, they were numbered 1–8, as shown in fig. 8, but the order in succession of horizons 3–8 had to remain tentative. it was however already clear from the older collections that there were more to be differentiated, given the necessary stratigraphical resolution. as already recounted, this resolution became available in the years 1994–1996, almost wholly in the region of ugleelv, and the number of faunal horizons in these zones has risen to thirty-four (fig. 30). they have therefore been given a new system of numbering, with prefixes bo-, inand poindicating the standard zones in which they lie (borealis, indistinctus and pompeckji). note that uncertainties in the faunal succession in the pompeckji zone discussed by callomon (1993, p. 96) have been eliminated, and the c. pompeckji, c. furcatus, c. carlsbergensis and c. gracilis horizons are shown to occur in reverse order compared to the 1993-scheme. the sections and levels in which the 34 faunal horizons have been recognised are shown in fig. 31. they show large gaps, but this must not be taken to mean that there are no sediments of the relevant ages in the successions. they are biostratigraphical gaps, meaning that the missing organisms were never present at the time or, if present, they left no preserved remains or reflect collectionfailure. to avoid having to renumber the faunal horizons above the cranocephalites beds and the pompeckji standard zone, which would be a cause for confusion, these higher horizons retain their old numbering unchanged, starting with horizon 9 at the base of the arcticus standard zone. they should be prefixed by the letter j (for jameson land) to avoid uncertainty (see fig. 8). standard zone stratigraphy the zonal stratigraphy of the cranocephalites beds up to 1993 was encompassed by the borealis, indistinctus and pompec kji standard zones, based on the then recognised succession of just eight faunal horizons. of these, five were included in an undifferentiated pompeckji standard zone. these zones form part of the secondary standard zonation for the boreal province of the biogeographic boreal realm of the middle jurassic (callomon 1959, 1993; see above). with the refinement available today, it is possible and useful to subdivide the pompeckji standard zone with its 23 45 46 callomon (1993) calyx variabile cranocephaloide ishmae greenlandicus c ar lsb er ge ns is po m pe ck ji in di st in ct us bo re ali s in te rm iss us g ra cil is ep isc op ali s arcticus pompeckji indistinctus 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 po-23 c. sp. (u-10) (u-9) (u-7) (u-6) (u-5) (u-4) (u-3) (u-2) (u-1) po-22 po-21 po-20 po-19 po-18 po-17 po-16 po-15 po-14 po-13 po-12 po-11 po-10 po-9 po-8 po-7 po-6 po-5 po-4 po-3 po-2 po-1 in-8 in-7 in-6 in-5 in-4 in-3 in-2 in-1 bo-3 bo-2 bo-1 c. episcopalis sp. nov. β c. episcopalis sp. nov. α c. ornatus (spath) c. maculatus spath c. transitorius (spath) β c. transitorius (spath) α c. sp. aff. gracilis spath c. gracilis spath c. aff. tvaerdalensis alsen c. tvaerdalensis alsen c. carlsbergensis callomon γ c. carlsbergensis callomon β c. carlsbergensis callomon α c. furcatus spath c. pompeckji (madsen) c. sp. aff. intermissus c. sp. c. sp. c. intermissus sp. nov. β c. intermissus sp. nov. α c. carolae sp. nov. β c. carolae sp. nov. α c. indistinctus callomon γ c. sp. c. indistinctus callomon β c. ex. gr. indistinctus c. indistinctus callomon α’ c. indistinctus callomon α c. sp. c. sp. c. borealis (spath) γ c. borealis (spath) β c. borealis (spath) α c.: cranocephalites spath 1932 borealis 1 this study fig. 30. the ammonite biohorizons of the borealis– pompeckji zones and the proposed subzonal subdivision of these zones compared to the zonation of callomon (1993). note that faunal horizons 1–23 of callomon (1993) were subsequently dubbed j-1 – j-23 (see fig. 8) to differentiate the jameson land zonation from that of milne land. 47 fo ss ilb je rg et f m pe lio n fm u nc er ta in ty fa un al ho riz on re co rd ed in a se ct io n po -2 3 e3 u gle el v se n w e2 e1 d 10 d 9 d 8 d 7 d 6 d 5 d 4 d 3 d 2 d 1 c 4 b4 b3 b2 b1 a c 3 c 2 c 1 po -2 2 po -2 1 po -2 0 po -1 9 po -1 8 po -1 7 po -1 6 po -1 5 po -1 4 po -1 3 po -1 2 po -1 1 po -1 0 po -9 po -8 po -7 po -6 po -5 po -4 po -3 po -2 po -1 in -8 in -7 in -6 in -5 in -4 in -3 in -2 in -1 bo -3 bo -2 bo -1 borealis zone indistinctus zone intermissus subzone carlsbergensis subzone gracilis subzone episcopalis subzone pompeckji zone fi g 31 .t he am m on ite b io ho riz on s o f t he b or ea lis –p om pe ck ji zo ne s a nd th e s ec tio ns in th e u gl ee lv ar ea in w hi ch th ey h av e b ee n re co gn ise d. faunal horizons into four subzones as shown in fig. 30. they reflect four successive basic morphologies of cranocephalites that should be recognisable more widely, making possible correlations at subzonal level of precision even when assignment to individual horizons is not possible. 48 taxonomy: general principles the middle jurassic ammonites of east greenland belong predominantly to only two families, the cardioceratidae and the kosmoceratidae. there were very minor, shortlived incursions of only two others, the oppeliidae (oxycerites) in the arcticoceras beds (middle bathonian) and the perisphinctidae (proplanulites) in the higher cadoceras–kepplerites beds (lower callovian). those to be described here, from the cranocephalites beds, all belong to the cardioceratidae and represent the earliest stages in the evolution of that family, which ranges from the late bajocian to the kimmeridgian (callomon 1985a). the fossil record of these ammonites in greenland is highly discontinuous, both in time and in space. the faunal succession has been pieced together from many localities and sections and synthesised in the time-ordered sequence of their distinguishable faunal horizons. the material from most of the horizons is abundant and in most cases well-preserved. in planning the taxonomic scheme to be adopted in its description, therefore, the following basic observations have to be taken into account: (1) from the definition of faunal horizons, it follows that the assemblage of fossils from any one of them must be regarded as isochronous on the time scale of what makes successive assemblages distinguishable: evolutionary change in their morphologies. fossil horizons may of course be locally ‘condensed’, i.e. contain mixtures of assemblages that are elsewhere biostratigraphically resolvable, but such condensation can only be recognised through fresh evidence from other localities at which it can be resolved. (2) the assemblages from single horizons consist overwhelmingly of adults. they show well the characters that are generally of great taxonomic importance in almost all ammonites. (3) the assemblages each consist of individuals whose morphologies, as far as one can tell, intergrade continuously. the test lies in adding further material, the effect of which is to make the spectrum of morphological variability trend increasingly toward unimodal, rather than accentuating polymodality. (4) the assemblages may therefore be regarded as monospecific. individual specimens are therefore treated merely as variants within a single, possibly highly variable biospecies. (5) successive faunal horizons yield assemblages that overlap considerably in their ranges of variability, i.e. share some morphological variants. the ‘vertical’ ranges of individual morphological variants, which were often in the past made into separate (morpho) species, can therefore extend over several faunal horizons and be members of several successive ‘horizontal’ biospecies. (6) such morphological overlapping is observed to greater or lesser degree over the whole range of the 34 faunal horizons. their biospecies are therefore linked as successive members of a single evolving lineage, the successive slices – transients – of a trunk of a family tree. such transients (a term introduced by bather in 1927) are sometimes referred to as chronospecies or chronosubspecies, but these terms are vulnerable to misinterpretation. the term chronospecies has also been used to label segments of a lineage ranging over several transients that the classifier regards are being still sufficiently similar to each other to be included in the same linnéan taxon. the boundaries between such successive chronospecies then become entirely subjective. subspecies are well-defined categories in the context of neontology that are used to label geographically restricted subsets of individuals within a common gene pool. although phenotypically distinguishable, they retain the potential to interbreed. such a ‘horizontal’, biospecific use of the subspecific category can in principle be validly retained in palaeontology if used in a similarly biogeographic sense to label isochronous local fossil races. but such purely biogeographic differentiation is rarely demonstrable in fossil assemblages. two distinguishable fossil assemblages at two distant localities may differ either because of isochronous genetic, racial differentiation, or because of heterochronous phyletic differentiation, or both. it is rarely if ever possible to decide between them. in an attempt to adopt a 49 natural, phylogenetic classification of a fossil succession, it is therefore best to avoid the use of the subspecific category altogether. a phyletic transient may then be labelled by continuing to use the name of an existing nominal species if, and only if, it can be shown that the type-specimen of the species came from the faunal horizon of that transient, i.e. that the horizon is the type horizon. other transients may be labelled by means of informal, additional nonlinnéan symbols such as α, β, γ … etc., rather than linnéan subspecific names. transients that differ sufficiently from those preceding them may be given new full specific names provided that it is understood that they refer to strictly isochronous ‘horizontal’ biospecies. the infra-subspecific category of variety, ‘var.’, should be retained in its strict sense as applied to variants within a biospecies. a suitable taxonomic category to encompass the succes sive transients of a monophyletic clade – the segments of a lineage – would then be the one of next higher rank in the linnéan hierarchy, that of genus. conventional classifications of ammonites have in the past used the category of the genus-group in a variety of ways. they shared, however, a common feature, again that of perceived purely morphological similarities in the form of selected shared characters. these similarities were now postulated to exist between ‘species’, but as these species were themselves purely morphological constructs, so were the morphogenera based on them. classifications oscillated between those of ‘splitters’ and those of ’lumpers’ and differed in the selection of the characters regarded as significant. the proliferation of new nominal taxa generated by the splitters cannot, under the taxonomic rules, be simply annihilated. they continue to swell the lists of synonyms of revisers even if suppressed as invalid taxa. but apart from the burden they impose on the literature, they can and do seriously mislead non-specialist taxonumerologists who take them at their face-value in counts to quantify species diversities and their changes with time in attempts to map the tempo of evolution and to identify its causes. thankfully, the number of nominal morphogenera that have been coined for the arctic pre-callovian cardio ceratidae is small. it has become clear that with only minor possible exceptions, the 50 or so successive transients of the cardioceratidae now recognised in the pre-callovian of greenland form but a single phyletic strand. it would therefore be possible, following the principles outlined above, to assign them all to but a single genus cadoceras, the oldest available name. but the lineage underwent some very considerable morphological transformations in the course of its evolution over this time-span. these changes have in the past been expressed through the use of the four successive genera of spath (1932) and used by him to label the four-way major biostratigraphical subdivision cited previously: cranocephalites–arctocephalites–arcticoceras– cadoceras. this usage has become so entrenched internationally that to abandon it now would be confusing. it continues to be useful because of the rough chronostratigraphical connotations that it implies, and it will therefore be retained here. but this is on the understanding that the dividing-lines between these genera are subjective and drawn strictly on grounds of convention and convenience. they do not mark phyletic discontinuities or phylogenetic bifurcation points or major morphological breaks. the faunas of the cranocephalites beds could in fact all be accommodated in the single genus cranocephalites, but the upper boundary is taken to lie at the level at which its successor, arctocephalites, has been conventionally introduced. finally, at the level of the family group, no special problems arise. the family cardioceratidae is used to accommodate a multiply-branching clade of lineages leading back to a single common origin. arising from this origin, the clade is subdivided both monophyletically and polyphyletically into several subfamilies. monophyletically, the main stem of the cardioceratidae has been conventionally subdivided into three successive segments, arctocephalitinae – cadoceratinae – cardioceratinae, to mark major changes of morphology, and this arrangement continues to be convenient. all the forms now to be described fall into the first of these and the subsequent developments of the others therefore need not be considered further here. the origin of the principal lineage of the cardioceratidae lies in fact in the earliest of the transients, cranocephalites borealis α, horizon bo-1 (callomon 1985a). its appearance in east greenland reflects an apparently abrupt re-colonisation by ammonites, after a long absence, of a large circumpolar boreal sea covering much of what is now northern siberia, northern alaska, arctic canada and the barents shelf, with an arm southwards to central east greenland and the shetland islands (see callomon 1985a, text-fig. 6a). the antecedents of the lineage lie most probably in the sphaeroceratid subgenus defonticeras of the upper lower bajocian of the northern pacific (callomon 1985a, textfig. 7). the main lineage of the cardioceratidae can then be followed at similar levels of time-resolution upwards to its abrupt termination in the kimmeridgian (callomon 1985a, text-fig. 3), in over 100 transients spanning some 20–25 ma. its habitat was exclusively boreal into at least the early callovian but then expanded somewhat southwards in what arkell (1956, p. 610) called the ‘boreal spread’. the first indications of a phylogenetic division of the main stem of the cardioceratidae point to levels no lower than the cranocephaloide standard zone (horizon j-18, fig. 8) and are indirect, expressed in the appearance at the base of the callovian of the genus chamoussetia (ch. menzeli mönnig 1995, p. 55), if this is taken to be derived from arcticoceras ishmae, horizon j-17 (callomon & wright 1989, text-figs 5, 6). the first positive evidence of a split lies in the observed co-occurrence of two biospecies (sensu callomon 1985a) of cadoceratinae in the nordenskjoeldi standard zone (horizon j-29 in fig. 8). others will doubtlessly emerge in the future. all the members of the succession of 34 faunas described here form with little doubt a single phyletic strand. 50 explanation and abbreviations collections, specimen abbreviations. specimens are all housed at the natural history museum of denmark in the geological museum section in copenhagen. four-digit numbers, prefixed jhc, refer to a catalogue by the first author of material largely collected by him in the years 1957–1996 in association at times with the late tove birkelund and the assistance of numerous helpers named in the acknowledgements. mguh are registration numbers at the geological museum, university of copenhagen. six-figure numbers, pre-fixed ggu, are serial field numbers from the geological survey of greenland (now merged with the geological survey of denmark into the geological survey of denmark and greenland, geus) that were assigned to material collected by various other workers, largely during the course of mapping in the years 1968–1974 by geologists from the university of copenhagen. specimen numbers prefixed casp were collected by geologists of casp, cambridge, uk; specimen numbers prefixed fs were collected by f. surlyk. types.ht: holotype; pt: paratype; lt: lectotype; [m], [m]: macroand microconch dimorphs respectively; at: allotype, secondary type, the opposite dimorph – the antidimorph – of the primary type. as the taxonomic dif ferentiations are based on comparisons of assemblages thought to represent isochronous biospecies, the descriptions take into account not only the characters of type specimens but also the variabilities of biospecies as seen in such assemblages. one of these assemblages of specimens found in association in a single bed at a single locality has therefore been selected as a reference collection in each of those species or transients in which material is sufficiently abundant. in most cases, the reference collection includes the types of the species or the representative specimen(s) of the transient it represents. paratypes have in consequence not generally been explicitly designated as such. w h u d d m ax d ph a b fig. 32. a: sketch of an adult/mature ammo nite (upon photo of c. pompeckji, see plate 11, fig.1a) illustrating the quantitative parameters of largest whorl diameter, dmax, and diameter dph at the last septum, indicated by the small arrow. b: sketch of an ammonite cross-section showing the following parameters: h: whorlheight. w: whorl-width. u: umbilical width, at diameter d. systematic taxonomy 51 dimensions. dmax: maximum diameter of adult shell at the peristome; dph: diameter at the last septum (marked with arrow) of the phragmocone (fig. 32a). h, w, u: whorlheight, whorl-width and umbilical width at diameter d (fig. 32b); their coefficients h (= h/d), w (= w/d), u (= u/d); < >: mean value in a sample of n measurements; σ: their standard deviation from the mean in an assumed simple gaussian distribution; ρ(π): the logarithmic spiral halfwhorl constant, the ratio of shell diameters half a whorl (π) apart. taxonomic descriptions. the taxonomic descriptions within the cranocephalites genus that follow are arranged in the temporal order of transients shown in fig. 30, grouped into standard zones and subzones and referred to their respective faunal horizons; the sources of the material are shown in fig. 31. the quality and quantity of the material varies widely and in a number of cases is insufficient to justify the creation of a new formal nominal taxon. such cases are included because they are nevertheless stratigraphically and morphologically sufficiently distinct from their neighbours to show that perceptible evolutionary change has occurred. they are distinguishable transients and are named in open nomenclature and assigned informal labels u-1 – u-10 (‘u’ for ugleelv). order ammonoidea suborder ammonitina superfamily stephanoceratoidea neumayr 1875 the names of taxa at superfamily level were not regulated under the international code of zoological nomenclature, until the advent of the current, fourth edition, whose provisions were effective as from january 2000 (international commission on zoological nomenclature 1999, 2012). article 29.2 now prescribes that names of superfamilies shall end in the suffix -oidea, whereas in previous editions such usage was put forward only as a recommendation (recommendation 29a). in what had become well-established convention in the taxonomy of ammonites and nautiloids, the treatise (treatise on invertebrate paleontology, moore 1957, 1964) had the names of superfamilies ending in -aceae, and this usage was continued in the second edition of the ammonite volume of the treatise dealing with cretaceous taxa (wright et al. 1996), whereas the recent contribution by howarth (2013) follows article 29.2. family cardioceratidae siemiradzki 1891 subfamily arctocephalitinae meledina 1968 genus cranocephalites spath 1932 type species. c. vulgaris spath 1932 [incl. boreiocephalites meledina 1967, type species b. pseudoborealis and c. (pachycephalites) meledina 1973, type species c. (p.) spathi]. the generic classification is discussed above. one generic name is regarded here as sufficient, but those wishing to distinguish the earliest forms, c. borealis, may retain boreio cephalites, perhaps as a subgenus. morphological characters of general taxonomic value. it is a general feature of almost all jurassic ammonitina that they grew to an adult stage at which growth stopped. experience with the greenland material has shown that the most useful characters for a natural classification of the arcto ce phalitinae are those measured at the adult stage: the diameter and whorl-section of the fully grown phragmocone, the modifications in coiling, length and sculpture of the mature body chamber, the form of the final peristome and the size of the fully grown adult. other morphological characters such as whorl-height, whorl-width and umbilical width have been found to be of lesser significance for they are liable to vary widely within a species. growth of the shells was logarithmically isometric in the planispiral diameter d, in the area of cross-section of the whorls a, and hence more or less in the coupled biometric parameters of whorl-height (h), whorl-width (w) and umbilical width (u) – the parameters usually cited – up to the end of the adult phragmocone. onset of the adult body chamber then leads to significant departures from this type of simple growth. conversely, the observation of such departures is diagnostic of the adult stage. common features are contraction of the whorl cross-sectional area that is reflected in a number of ways: (1) an uncoiling of the umbilical seam; (2) contraction of the interseptal spacing reflected in crowdingtogether (approximation) of the last septal sutures; (3) modification of the sculpture of the shell, commonly by the loss of all ribbing (variocostation); and (4) a change in the sweep and structure of the final peristome, in the arctocephalitinae commonly in the form of a terminal constriction on internal moulds, reflecting a terminal strengthening collar on the shell. the average final adult diameter of a transient assemblage is a quite closely-defined character, particularly significant in the recognition of sexual dimorphism. because of the variable modifications of the adult body chamber observed in this material, biometric measurements of the principal parameters of the shell morphology at the maximum shell-diameters have little comparative 52 value. values of h, w, u, where given, are therefore quoted generally at the end of the adult phragmocone, at dph. in most transients, the values of h and u vary so little between subsequent transients that they are hardly worth measuring. the whorl-width, w, does, however, vary intraspecifically and from transient to transient, so it is generally given. ranges of intraspecific variation are indicated in terms of a gaussian model, standard deviation σ. all these features are prominently well developed in the faunas described below. the genus cranocephalites is strongly dimorphic. the adult macroconchs are all more or less variocostate, the ribbing modifying and tending to fade, with simple final peristomes preceded by a broad, shallow constriction on the internal moulds. the microconchs are isocostate, with a simple but ventrally projected peristome. the dimorphic size ratio is large, in the range 2.5:1 to 3:1, indicating that the body-chamber volume of macroconchs is much larger than in microconchs. palaeoecology. almost all of the abundant assemblages of ammonites collected in east greenland consist almost exclusively of fully grown adults and predominantly of macroconchs at that. microconchs are rare, despite the fact that a special watch was kept out for them in the field. the sex ratio of [m]:[m] is therefore in the range of 10:1 – 100:1. in some of the assemblages to be described, no microconchs have been found at all. the probable reasons for these ontogenetic and sexual polarisations have been discussed previously (callomon 1985a). they point to a closed life cycle in which the fossil accumulations now mark the breeding, spawning and dying grounds of gregarious nektonic organisms that migrated during their life cycles, with sexual segregation in the final stages. the persistence of such characteristic death-assemblages of mature adults of one genus in a geographically restricted region for such long periods of time, as seen in jameson land, points to that region as the ‘true home’ of that genus for reproductive purposes and hence for its evolution. such assemblages have been called eudemic and those of east greenland taken as prototypical (callomon 1985a). juvenile ammonites do occur but in quite different surroundings. the otherwise unfossiliferous, more distal, finegrained mudstone-dominated successions occasionally yield an isolated concretion packed with completely-preserved small juveniles, too young to identify even as [m] or [m]. the impression is that of a death-assemblage of a migrating school of gregarious juveniles. borealis standard zone the borealis standard zone includes three faunal horizons bo-1 – bo-3, characterised by transients of cranocephalites borealis (spath 1932; fig. 30), and illustrated in plates 1–3. three separable transients α, β, γ of these, the earliest forms of cranocephalites, can now be distinguished in jameson land. the type of the original species came most probably from trans α. the three transients differ relatively little and some variants are found through the vertical range of all the transients. conversely, a single specimen found in isolation may not be assignable to any of the transients in particular. the features they have in common are as follows: general description. strongly dimorphic. the macroconchs are on the whole smaller than those of younger/succeeding species. they also differ in that maximum diameters within an assemblage can vary much more widely. mature body chambers occupy c. 0.75 whorl and contract strongly, the umbilical seam uncoiling markedly. the final peristome is simple, preceded by a prominent, broad but shallow constriction on the internal mould. the coiling is strongly involute, with minute umbilici on the inner and middle whorls. the whorl-section can range from slightly high oval in compressed variants to rounded, depressed in inflated variants. the umbilical shoulder is always well rounded. the most characteristic feature lies however in the style of the ribbing: coarse but blunt and subdued on the phragmocone, primaries rising gently on the umbilical shoulder, dividing without accentuation into equally subdued secondaries interspersed with intercalatories that cross the venter without loss of strength. the ribbing tends to fade on the adult body chamber, which becomes wholly smooth in large specimens. the microconchs are small but otherwise resemble the macroconchs in every respect. their adult sizes are again very variable and a large microconch may be hard to distinguish from the smallest macroconch. the microconchs remain ribbed to the end, however. the adult stage is seen in the uncoiling of the umbilical seam and in the whorlwidth of the body chamber, which ceased to grow isometrically and stayed constant. the peristome is also simple and preceeded by a constriction that is however shallower and broader than in the macroconchs. it is these expressions of the dimorphism that reveal that the origin of cranocephalites, and hence of the whole of the cardio ceratidae, lies in late early bajocian sphaeroceratidae of the east pacific realm (callomon 1985a). 53 bo-1:cranocephalites borealis (spath 1932) sensu stricto, trans α plate 1, figs 1–5 [m], figs 6–7 [m]; fig. 33a, b 1932 xenocephalites borealis spath, p. 44, plate 14, fig. 4a–d (ht by monotypy). 1984 cranocephalites borealis (spath) [m] – callomon, p. 148, fig. 2b [jhc 1368]. 1985a cranocephalites borealis (spath) [m] – callomon, p. 64, fig. 7b [id.]. type specimen. the precise level within the range of bo-1 – bo-3 from which it came is somewhat uncertain. its morphology is of little help, for it is a wholly septate macroconch nucleus only 24 mm in diameter (fig. 33a, b). it is, however, very strongly and coarsely ribbed, suggesting the forms that retain their strong ribbing longest, which are those of trans α. but similar nuclei are equally common in bo-2, trans β. the place of origin is not much help either. it was collected by rosenkrantz at ‘mt hjørnefjæld’ (hjørnefjeldet, fig. 4) allegedly at a height of 630 m, which, according to the sketch of his section (in spath 1932, p. 133, text-fig. 14), places it on the slope of the sortehat formation. the only associated specimen, a crancocephalites cf. furcatus (? po-9), definitely came from the local ‘cranocephalites bed’, at 740 m in the section. there is in fact a further complication. it relates to the precise location of ‘mt hjørnefjæld’ itself. rosenkrantz’s sketch-maps (in spath 1932, p. 125, text-fig. 9, and rosenkrantz 1934, plate 1) places it at the ‘corner’ between the headwaters of depotelv, running westwards, and lejrelv, coming from north to south. but here the highest point lies at only 630 m, nothing like the 765 m given in the section. the nearest place corresponding stratigraphically to the section reproduced by spath lies at centralbjerg, sections 35–37 in fig. 1b, about 3 km west of the locality now marked as hjørnefjeldet on the 1:100 000 geological map of the geological survey of greenland (ggu) (birkelund & higgins 1980). its highest point lies at 660 m, and the succession there can be brought roughly into correspondence with rosenkrantz’s section if all the heights given in the latter are reduced by 100 m. rosenkrantz and co-workers used aneroid barometers to measure heights, and differences in height indications among authors may thus just reflect differences in atmospheric pressure. the highest beds on centralbjerg are in typical pelion formation and are indeed only a little above a prominent local marker, rosenkrantz’s ‘arctocephalitesbed’, rediscovered in 1971 and now recognised as horizon j-9 (fig. 8). rosenkrantz also picked up a second prominent local marker which he called the ‘pecten-belemnite horizon’ (bed (d) in his section, shown in spath 1932, p. 133 fig. 14), 60 m below the arctocephalites bed and 50 m above the top of the sortehat formation. spath referred to it as the demissusgrit (after entolium demissum (phillips), rosenkrantz’s ‘pecten’) and surmised that it was the source of the type of c. borealis. this marker bed was also rediscovered in 1971 during mapping and referred to in the field as the ‘brown bed’, c. 0.5–1 m thick, for its colour makes it stand out in the otherwise monotonous succession of white or cream-coloured sandstones of the pelion formation within which it lies. also, in contrast to the beds above and below, it is fossiliferous, with abundant belemnites, some oysters and other bivalves. the bed can be seen all round centralbjerg and followed across to the northern slopes of mikael bjerg (section 33, bed 4, 6 km wsw of section 35, fig. 4.). at centralbjerg it lies about 100 m above the sortehat formation and 100 m below the arctocephalites bed. at mikael bjerg, it was estimated to lie about 90 m above the sortehat formation and 80 m below the arctocephalites bed; at this locality, it yielded four spe cimens of c. borealis α. the ‘brown bed’ is the only bed in the pelion formation in the mikael bjerg – hjørnefjeldet area to have yielded ammonites below the pompeckji standard zone. the conclusion has therefore to be that the type-horizon of c. borealis most probably was the ‘pecten-belemnite horizon/brown bed’ at or close to cen tralbjerg and that it is the lowest of the three ammonite horife b a c dc ddd 1 cm fig. 33. a, b: the holotype of cranocephalites borealis [m], mguh 9235. c, d: ‘xenocephalites’ borealis, nordenskjöld collection 1900, near fossilbjerget. e, f: topotype of c. warreni frebold 1961 from richardson mountains, yukon, canada (jhc collection 1975). all natural size. 54 zons with c. borealis, bo-1, trans α. it seems ironical that the type that was brought back after the extensive surveys of 1926–1927 should have been the sole specimen, and it from the rarest transient, of what is in fact one of the most abundant and widespread species in the whole of jameson land. there was in fact a further specimen that had been brought back by nordenskjöld in 1900 during the danish amdrup expedition, from his ‘fossil locality 1’, somewhere in the region of fossil mountain. it was referred to by madsen (1904, p. 198, no. 5) but not otherwise described. spath saw it (1932, p. 45) and gave something of a description, assigning it to his ‘xenocephalites’ borealis, but did not figure it. he noted its small size and that it carried threequarters of a whorl of body chamber, with rejuvenation of the ribbing at the peristome. it is now figured here, in fig. 33c, d. although poorly preserved, it is clearly an adult microconch. its preservation is in brown sandstone similar to that of the ‘brown bed’ and may well have come from it, hence also trans α. in view of the inadequacies of the formal type specimen, the one shown on plate 1, fig. 3a, b (jhc 1368) may be taken as typical stand-in (typus substituens) for the macroconchs, and that on plate 1, fig. 7, (jhc 1363 [m]) as allotype. material and distribution. ugleelv, section e2 (reference collection), bed 12, top of ug-1, 80 m above sortehat formation (fig. 28a): 30 [m], 3 [m]. ugleelv, section d10 (fig. 26): 2 [m], 3 [m]. mikael bjerg, section 33, bed 4, ‘brown bed’, c. 110 m above base of pelion formation: 2 [m], 2 [m]. trefjord bjerg, section 39 (fig. 29): 3, loose, from the basal 20 m of the more massive sandstones of the pelion formation (see further discussion below). total: 44 (37 [m] + 7 [m]). description. the macroconchs range in size from 40 to 70 mm (plate 1, figs 4, 5), the microconchs from 20 to 35 mm (plate 1, fig. 7). the adult body chamber becomes constant in whorl-width (plate 1, figs 1–3, 6). the whorl-width varies widely, from 0.6 (plate 1, fig. 2) to 0.8 (plate 1, fig. 1) of the diameter at the last septum. the shells retain their characteristic strong ribbing to the end, with at times even some strengthening at the final stage (plate 1, fig. 5). dimensions. discussion. the species was initially placed in the genus xenocephalites spath 1928, because of all the middle jurassic ammonites then known it was the only one whose type species remotely resembled the greenland specimen. the figure of the type of the type species, macrocephalites neuquenensis stehn 1924, is reproduced in the treatise (arkell et al. 1957, fig. 352.13). the interpretation of the genus has since become clear. it represents the microconchs of another sphaeroceratid group, the eurycephalitinae that is endemic in the east pacific realm, extending from alaska to the andes and thence to the palaeo-moluccas of indonesia (riccardi & westermann 1991). bo-2: cranocephalites borealis (spath 1932) trans β [m]: plate 1, figs 11–13; plate 2, figs 1–10 [m]: plate 1, figs 8–10; plate 2, figs 11–14 cf 1957 xenocephalites kononovaja voronets, p. 22, plate 1, fig. 3a, b, v (ht by monotypy). cf 1957 morrisiceras laptinskajanavoronets, p. 21, plate 1, fig. 2a, b, v (ht by monotypy, [m]?). ?1957 morrisiceras sibirica voronets, p. 22, plate 1, fig. 1a, b, v, g (ht by monotypy). 1959 cranocephalites borealis (spath) – callomon, p. 507, plate 17, figs 1a, b, 2a, b; plate 18, fig. 3a, b. 1961 cranocephalites borealis (spath) – frebold, p. 12, plate 1, figs 1–4. cf 1961 cranocephalites warreni frebold, p. 14, plate 2, figs 1, 2, 4. 1962 xenocephalites kononovaevoronets, p. 45, plate 9, fig. 3 (= 1957 redescribed). ?1962 morrisiceras laptinskajae voronets, p. 30, plate 15, fig. 3a, b (= 1957 redescribed). ?1962 morrisiceras sibiricum voronets, p. 31, plate 15, fig. 2a, b, v (= 1957 redescribed). 1967 boreiocephalites pseudoborealis meledina, p. 107, plate 1, figs 1a, b, v, 3a (ht), figs 2a–g, 3b (pt). dmax dph h w u jhc 1368 (plate 1, fig. 3): 57 39 0.50 0.74 0.12 jhc 1789: 53 37 0.50 0.58 0.12 mean values, reference collection: <dmax> 51.7 ± 1.9 σ = 8.4 (16%) n = 21 <dph> 31.6 ± 1.8 σ = 9.1 (29%) n = 26 <wph> 0.66 ± 0.02 σ = 0.11 (16%) n = 26 dmax, dph in mm 55 1973 boreiocephalites pseudoborealis meledina, p. 35, plate 1, figs 1, 3 (same as 1967), 2a, b. 1982 cranocephalites borealis (spath) – poulton et al., p. 84. 1984 cranocephalites borealis (spath) [m] – callomon, p. 149, fig. 2b. 1985a cranocephalites borealis (spath) – callomon, p. 64, text-figs 7b, 8a; [m]: text-fig. 8a. 1988 boreiocephalites borealis (spath) – krymholts et al., p. 22, plate 5, fig. 7a, b (=meledina 1973, plate 1, fig. 1). representative specimen of transient β [m]. mguh 22255 (ex jhc 1144), plate 1, fig. 12 (= callomon 1959, plate 17, fig. 1a, b); in the reference collection from teebjerg, section c1, bed 18. representative specimen of transient β [m]. ggu 135887a, plate 2, fig. 14; katedralen, section d9, bed 9. material and distribution. this is one of the most abundant and widespread forms in the ugleelv area. it occurs at several levels in the ug-2 unit but the abundance at any one of these can change rapidly from place to place. by far the richest occurrence was in a single thin level of slightly phosphatised concretions at teebjerg, section c1, bed 18, the reference collection, that littered the ground in their hundreds where weathered out, many of them enclosing a beautifully preserved ammonite. yet at falkeelv, section c4, 2 km south-west and at statuebjerg, section b3, 2 km east (fig. 9), the bed had disappeared. reference collection: section c1: 185 [m], 3 [m]. section b8: 8 [m]. section c4: 2 [m]. section d8: 6 [m] (fig. 25). section d9: 78 [m], 11 [m]. section d10 (fig. 26): 29 [m]. section e1: 6 [m] (fig. 27). section e2 (fig. 28): 11 [m]. total: 339 (325 [m] + 14 [m]). poorly preserved material was often noted but not collected. an important example was on the north ridge of trefjord bjerg, section 39b (fig. 29), where it occurred irrecoverably in beds 5–10 in the pelion formation, 80 m above its base and immediately below the main cranocephalites horizon there, that of c. carlsbergensis, fauna po-11. description. the range of variability is shown in the plates. in contrast to trans α (plate 1, figs 1–4), the adult macroconchs of trans β become smooth on the body chamber (plate 1, figs 11–13). the maximum sizes range from 37 to 83 mm (plate 2, fig. 2). the whorl-width can range from 0.50 of the diameter (plate 2, fig. 2) to over 0.90 (plate 2, figs 1, 10). the whorl-section can be ventrally rounded (plate 2, figs 6, 8, 10) or arched (plate 2, fig. 4). the characteristic terminal constriction is well seen on the representative specimen (plate 1, fig. 12). a slightly immature or precociously aged specimen is shown in plate 2, fig. 5: the ribbing persists, the body chamber is somewhat longer than usual (0.85 of a whorl, vs. 0.75) and the peristome has not yet modified. a typical nucleus is shown in plate 2, fig. 3. microconchs range in size from c. 15 to 30 mm (plate 1, figs 8–10; plate 2, figs 11–14). their peristomes have typically only a narrow, inconspicuous constriction followed by a slight lateral flaring, well seen in plate 2, figs 11 and 13a, b. dimensions. discussion. the synonymy includes specimens from the richardson mountains in the yukon territory (frebold 1961) and siberia (voronets 1962; meledina 1968) described under various new names. the material consists in each case of but a few specimens, insufficient to define an assemblage with a precision comparable to that of those being described here. the specimens match as far as they go the greenland forms of trans β of comparable sizes, their body chambers becoming wholly smooth and the ribbing exactly as in c. borealis. a topotype of c. warreni from the yukon collected in 1975 (poulton & callomon 1976; poulton 1978) is figured here (fig. 33e–f). bo-3: cranocephalites borealis (spath 1932) trans γ plate 3 representative specimen of trans γ. jhc 1011 (plate 3, fig. 6a–c), in the reference collection from hurry inlet, section 8, brinkmann fjeld, bed 3 (fig. 2). material and distribution. the pelion formation unit ug2 around ugleelv terminates everywhere with soft sandstones containing a level of small round concretions carrying large plagiostoma bivalves, topped by thin lenticular red ironstones dmax dph h w u mguh 22255: 58 ~41 at d = 40 0.50 0.61 0.10 (0.75 whorl body chamber) mean values, reference collection: <dmax> 52.0 ± 1.2 σ = 7.4 (14%) n = 36 <dph> 36.8 ± 0.9 σ = 6.1 (16%) n = 47 <wph> 0.625 ± 0.01 σ = 0.060 (9.5%) n = 47 d, dmax, dph in mm 56 and followed with sharp change of facies by the soft, silty mudstones of unit ug-3. the top few metres of ug-2 yield the last c. borealis, differing somewhat in aspect from those of trans β below and now differentiated as trans γ. in the region of hurry inlet, at neill klinter, ug-2 terminates in similar fashion but is immediately followed there by the mudstones of the arcticus standard zone, horizon j-10, and both the indistinctus standard zone and the pompeckji standard zone are missing. here, too, the top of ug-2 carries concretions with c. borealis γ. they are particularly abundant around moskusoksekløft at brinkmann fjeld, which produced the best-characterised assemblage, now taken to be the reference collection. found also at scattered localities elsewhere in jameson land while mapping. hurry inlet, sections 1–8: 34 [m]. ugleelv: sections b2 (fig. 19), c2, c4 (fig. 21), d2 (fig. 22), d5 (fig. 24), d9, e3: 52 [m]. total: 86 (all [m]). description. inner and middle whorls of macroconchs indistinguishable from those of trans β at comparable diameters, except that more inflated variants dominate. the most striking difference lies however in the size (plate 3). as the ribbing is lost at about the same diameters as in trans β, all that is generally visible in trans γ is entirely smooth. the microconchs have yet to be found but are expected to be very much like those of trans β, perhaps a little larger. dimensions. an estimate could be made of the logarithmic spiral halfwhorl constant from specimens that had been broken diametrically (cf. plate 3, fig. 6): ρ(π) = 1.29 (± 0.04), n = 4. the phyletic increases in the sizes of the adult shells of c. borealis are reflected in the mean values of the diameter of the phragmocones <dph> in successive transients listed above: α: 32 (± 2) → β: 37 (± 1) → γ: 50 (± 2) mm. general discussion of the cranocephalites borealis group other species. besides the faunas bo-1 – bo-3 there are indications of at least two other species of cranocephalites of the borealis group, one from east greenland, the other from siberia, each described in terms of two published nominal morphospecies. east greenland: (1) cranocephalites kochidonovan 1953, (p. 80, plate 15, fig. 8a, b [m], generically as subgenus of arctocephalites), from localities 120 and 104 on traill ø, including cranocephalites kochi var. latus dono van (p. 82, plate 16, fig. 6a, b,[m] from locality 137 and cranocephalites kochi var. pygmaeus donovan (p. 81, plate 16, fig. 5a, b [m?]) from localities 120, 137; (2) cranocephalites parvusdonovan 1953 (p. 79, plate 15, fig. 2a, b [m]), also from locality 120. stratigraphically, their positions are not very clearly defined. morphologically, however, they share their style of coiling and ribbing with those of c. borealis but are consistently smaller, none of them exceeding 45 mm in diameter. they resemble c. borealis α in retaining strong ribbing on the adult body chamber, but the ribbing on c. parvus is markedly denser and finer than on any of the transients α–γ of c. borealis, recalling perhaps more strongly the putative ancestors of the whole lineage, chondroceras of the early bajocian. although the amount of material from traill ø is small, it is enough to show that it differs consistently from c. borealis. the only known forms from jameson land that could match it are from trefjord bjerg, section 39 (fig. 4), from the basal 20 m of the pelion formation (ug1), hence presumed to be at least as old as bo-1, c. borealis α. only three specimens exist, two of them poorly preserved. they are also small, not exceeding 50 mm in diameter, and more densely and finely ribbed than the usual c. borealis α. one of them, although crushed, could in this respect be a close match with c. parvus. if these indications are correct, they make c. kochi/parvus the oldest known species of cranocephalites and the founder of the cardio ceratidae. siberia: (3) cranocephalites sibiricus (voronets 1957) (p. 22, plate 1, fig. 1a–g, originally as morrisiceras, reproduced and described more fully in voronets 1962, p. 31, etc.). dmax dph h w u jhc 1011: at d= 60 (0.70 whorl body chamber) mean values, reference collection: <dmax> 74.5 ± 3.4 σ = 15.7 (21%) n = 22 <dph> 49.7 ± 1.6 σ = 8.2 (16%) n = 27 <wph> 0.63 ± 0.02 σ = 0.12 (19%) n = 28 d, dmax, dph in mm 0.51 6185 0.85 0.13 57 (4) cranocephalites laptinskajae (voronets 1957) (voronets 1962, fig. 2a–v). these morphospecies, although still clearly related to c. borealis in coiling and the blunt ribbing, differ in retaining strong secondary ribbing long after the primaries have been lost, leaving a wholly smooth umbilical shoulder. both came from the ury ung –tumus peninsula, west of the anabar estuary in northern siberia, allegedly from ‘bed 5’. the section was reproduced by meledina (1973, p. 104) but the ammonites are recorded from ‘bed 6’ (16 m), together with cranocephalites spp. of the pompeckj–furcatus–maculatus group, clearly indicating the pompeckji standard zone. at least, beds 5 (21 m) and 6 are above bed 4 (5 m), which yielded c. borealis proper, boreiocephalites pseudoborealis meledina, identified with trans β above, so that c. sibiricusmust be younger than c. borealis. there would be room in the succession in greenland for additional faunal horizons at the regional facies change from ug-2 to ug-3 and additional faunal transients at the marked morphological change at that level, between bo-3 and in-1. world-wide distribution. arctic canada (yukon), northern siberia (anabar, lena), central east greenland. indistinctus standard zone the indistinctus standard zone includes eight faunal horizons in-1–in-8, represented by the group of cranocephalites indistinctus callomon 1959 and cranocephalites spp. (u-1) – (u-4) (fig. 30), illustrated in plates 4–8. the marked lithological break from ugleelv units ug-2 to ug-3 marks also a fundamental change in the morphology of cranocephalites. henceforth the shells are more densely and sharply ribbed on the phragmocones, the primaries and secondaries less differentiated, the variabilities of assemblages reduced and the features distinguishing successive transients expressed mainly in the form of the adult macroconch body chamber. the soft, recessive mudstones of ug-3 make the bio stratigraphy heavily dependent on the relatively uncommon, impersistent and often very poorly fossiliferous layers of doggers or indurated mudstone. the ammonite record in this interval remains patchy. although eight distinguishable horizons are now recognised, in only a few of them is the material sufficiently abundant and well preserved to be worth describing in any detail. four of them are labelled in open nomenclature with numerical symbols: (u-1) – (u-4). the type material of c. indistinctus came from scattered, isolated concretionary levels in ug-3 between statuebjerg (sections b1, b2; figs 18–19) and ræveelv (section e3) (callomon 1959). it was clear that more than one faunal horizon was represented, but even after new collecting in 1970–1971, only two could be recognised with any confidence, labelled c. indistinctus α and β (callomon 1993, p. 96, fig. 4). these can now be located in a more extensive succession. in-1, in-2: cranocephalites spp. (u-1), (u-2) (u-1): plate 4, fig 1a, b, [m] (u-2): plate 4, figs 2–4 [m], figs 5, 6, [m] material and distribution. (u-1): representative specimen jhc 5306 (plate 4, fig. 1, [m]), section d2, 6 m above the plagiostoma bed (fig. 22): 8 [m], 1 [m]. (u-2): representative specimen jhc 5314 (plate 4, fig. 2a, b, [m]), section d2, 10 m above the ‘plagiostoma bed’: 18 [m], 2 [m]. total: 29 (26 [m] + 3 [m]). in sections d2 and d5 (figs 22, 24), these specimens were found at two levels 3–4 m apart on the particularly featureless lowest 10–15 m of mudstones above the ug-2 sandstones, on gentle slopes that are usually obscured by scree. exceptional rainfall in 1994 had, however, provided clean exposures, degraded again already in 1996. the ammonites had originally been embedded freely in the mudstones, acting as their own concretions. most have now lost their lithified body chambers and all that remains are the partially crystal-infilled phragmocones. descriptions. macroconchs with subcircular whorl-section, slightly depressed (w/h = 1.23), with narrow, deep umbilici and rounded umbilical shoulders; ribbing dense, strong and sharp, dividing irregularly with intercalatories at whorl mid-flank, persisting to the end, only mildly variocostate; length of body chamber 0.75 whorl. (u-2) more strongly, sharply ribbed than (u-1), and the whorl-sides somewhat less rounded. the differences are at the limit of what can be morphologically resolved in assemblages of moderate size, especially when lacking the adult body chambers. two microconchs of (u-2) are shown in plate 4, figs 5, 6. they too are complete phragmocones but the uncoiling umbilical seams show the extent of the adult body chambers to be up to 0.75 of a whorl, as in the macroconchs. estimated complete diameters 32 and 40 mm, respectively, making the average dimorphic size-ratio only 2:1, unusually low for cranocephalites. 58 dimensions. discussion. the abrupt change in morphology from that of c. borealis is remarkable, particularly in the density and sharpness of the ribbing. the form of the phragmocone remains practically unchanged to the top of the pompeckji standard zone. inner whorls of many successive transients may be practically indistinguishable. in-3 – in-8: cranocephalites indistinctus callomon 1959 1959 cranocephalites indistinctus callomon, p. 510, plate 17, figs 3, 4; plate 18, figs 1a, b, 2a, b. ?1973 cranocephalites indistinctusmeledina, plate 4, fig. 3a, b. 1985a cranocephalites indistinctus [m] callomon, fig. 8b (allotype). 1993 cranocephalites indistinctus α, β callomon, p. 96. holotype. cranocephalites indistinctuscallomon, 1959, plate 18, fig. 1a, b, (jhc 1435). in-3: cranocephalites indistinctus trans α (sensu callomon 1993) plate 4, figs 8–13 [m]; figs 7, 14 [m]. 1959 cranocephalites indistinctus callomon, plate 17, fig. 4. 1985a cranocephalites indistinctus [m] callomon, fig. 8b (allotype). 1993 cranocephalites indistinctus α callomon, p. 96. representative specimen of transient α [m]. mguh 22258 (ex jhc 1306), plate 4, fig. 11a, b; section b3, bed 20. representative specimen of transient α [m]. jhc 1248 (plate 4, fig. 14); section b1, bed 24 (fig. 18). material and distribution. ugleelv sections: b1–b3: 9 [m], 3 [m]. section d2 (fig. 22): 23 [m], 7 [m]. section d10 (fig. 26): 1 [m]. total: 43 (33 [m] + 10 [m]). description. somewhat more consistently compressed, smaller, more feebly and densely ribbed to the end than the forms below or those in the pompeckji standard zone above. dimensions. in-4: cranocephalites indistinctus trans α´ plate 5, figs 1–3 [m] representative specimen of transient α´ [m]. jhc 6017 (plate 5, fig. 1). material. from a single locality, section b2, bed 12a, 5 m above in-3 (trans α), 7 m above the top of ug-2 (fig. 19): 6 [m]. description. the main features as in trans α, but differs in a more strongly variocostate adult body chamber, becoming smooth. umbilicus extremely narrow. dmax (u-1) dph h w u jhc 5306: at d= 40 (0.75 whorl body chamber) mean values, reference collection: <dmax> 82 (est.) σ = n = 10 <dph> 51 σ = n = 9 <wph> 0.63 ± 0.02 σ = 0.038 (6%) n = 8 dphdmax (u-2) h w u ρ(π) jhc 5314: wholly septate at d= 46 (0.75 whorl body chamber) mean values, reference collection: <dmax> 75 (est.) σ = n = 10 <dph> 46 σ = n = 10 <wph> 0.59 ± 0.02 σ = 0.063 (11%) n = 10 d, dmax, dph in mm d, dmax, dph in mm 0.51 4363 0.59 0.13 0.48 46 0.69 0.16 1.37 dmax [m] dph h w u mguh 22258 (ex jhc 1306): (0.8 whorl body chamber, crushed, with peristome) mean values (a collection of 20 specimens from a single concretion) <dmax> 0.54 ± 1.2 σ = 4.1 (8%) n = 13 <wph> 0.56 ± 0.00 σ = 0.035 (6.3%) n = 9 [m] jhc 1248: <dmax> = 24 mm, n = 3; other measurements not possible. dimorphic size-ratio 2.2 dmax, dph in mm 3760 59 dimensions. in-5: cranocephalites ex. gr. indistinctus (u-3) (not figured) material and distribution. another distinct assemblage from an isolated locality, section e1, bed 9 (fig. 27), a layer of sparse doggers in ug-3 mudstones 10 m above the top ironstone of ug-2: 6 [m], poorly preserved, crushed. description. the smallest of the group, ranging in size from 46 to 60 mm when fully grown (<dmax > = 50 mm (n = 4)), densely and finely ribbed to the end. the order of in-4 and in-5 in the succession is conjectural, as both are from single levels in the variably thick ug-3 mudstones from two localities 6 km apart. in-6: cranocephalites indistinctuscallomon 1993 sensu stricto, trans β plate 5, figs 4–6, plate 8, figs 1, 2 [m] 1959 cranocephalites indistinctus callomon, p. 510, plate 17, fig. 3; plate 18, figs 1a, b, (ht), 2a, b. 1985a cranocephalites indistinctus [m] callomon, text-fig. 8b (ht). 1993 cranocephalites indistinctus β callomon, p. 96. holotype [m]. mguh 22259 (ex jhc 1435; plate 5, fig. 4); section c2, bed 28. paratypes [m]. jhc 1438 (paratype i), plate 5, fig 5; jhc 1439 (paratype ii), plate 5, fig. 6. material and distribution. the type collection, section c2, 7 m above the top of ug-2: 17 [m]. section b1 (fig. 18), 18 m above top ug-2 and 2 m above c. indistinctus trans α: 8 [m]. section e2 (fig. 28), 16 m above the top of ug2, immediately below the base of the clinoform bed of ug4: 2 [m]. total: 27 [m], microconchs unknown. another example of the sporadic local concentration of these ammonite assemblages: so abundant at teebjerg, section c2, the same still clearly recognisable at statuebjerg n, section b1 (fig. 19), but no trace of either the bed or the ammonites at nearby section b2 (fig. 18). description. the macroconchs resemble those of trans α´, in-4 (cf. plate 5), but are larger, more strongly ribbed and with more open umbilicus through earlier egression of the umbilical seam at maturity. some variants are homeomorphs of c. gracilis in po-15. the range of variability may be seen in plate 8, fig. 1a, b, coarsely ribbed to the end; fig. 2a, b, finely ribbed, fading, compressed. dimensions. in-7: cranocephalites ex. gr. indistinctus (u-4) (not figured) material. two half-crushed macroconchs only (jhc 6023, 6024), from statuebjerg, section b2 (fig. 19), c. 12 m above ug-2, in a single concretion found on a gentle, clean slope of soft mudstones of ug-3, well separated from other hard beds, so that contamination by downward slippage from higher horizons can be safely ruled out. description. large, whorl-section well rounded, ribbing uniformly fairly dense to the end, fading on the adult body chamber but not disappearing altogether, rejuvenating towards the peristome. general overall resemblance in size and coiling is to the later faunas of po-1 – po-4, but in these the ribbing on the adult body chamber is stronger, coarser and more strongly variocostate. dmax dph h w u jhc 6017: (0.70 whorl body chamber) mean values: <dmax> 68 σ = n = 6 <dph> 48 σ = n = 6 <wph> 0.55 σ = n = 3 <uph> 0.05–0.09 σ = n = 2 dmax, dph in mm ~4568 dmax dph h w u mguh 22259 (ht): (0.75 whorl body chamber; other measurements not reliable. rather a small variant.) mean values, others: <dmax> 72 ± 0.9 σ = 6.5 (9%) n = 15 <dph> 48 ± 1.5 σ = 5.8 (12%) n = 16 <wph> 0.50 σ = n = 4 dmax, dph in mm ~4060 60 dimensions. in-8: cranocephalites indistinctus trans γ (not figured) representative specimens of transient γ. jhc 6031 [m]; jhc 6047 [m]. material and distribution. statuebjerg, section b2 (fig. 19), bed 13b, 16 m above ug-2, another clearly-defined horizon of concretionary sandstone well separated from others; ammonites abundant but mostly poorly preserved, fragmentary or crushed: 14 [m], 1 [m]. katedralen n canyon, section d8 (fig. 25), bed 15, a very similar sandstone 17 m above ug-2, a prominent marker in an otherwise featureless series of mudstones and sandstones with the long vertical burrows of diplocraterion habichi and containing abundant burrowing razor-shells, strimodiolus elongatus fürsich: 5 [m]. total: 20 (19 [m] + 1 [m]). description. a small species, but the maximum size is unusually variable, ranging from 48 to 75 mm; section sub-circular to depressed; densely and sharply ribbed to the end in small variants, fading on the body chambers of the large variants, which resemble the two large specimens of in-7. umbilicus slightly more open than in other transients. dimensions. general discussion. this is the youngest of the series of faunas typified by c. indistinctus: rather small, delicately ribbed to the end. it resembles overall the earlier forms, but the inclusion of larger forms presages the change to the subsequent dominant morphology of the faunas of the pom peckji standard zone. similar small forms of cranocephalites have also been found in the canadian arctic (prince patrick island: frebold 1958, plate 8, figs 2, 3, closest in resemblance to in-3, c. indistinctus α; richardson mountains, yukon: poulton et al. 1982, p. 84). the form described under the name c. indistinctus from the anabar region of northern siberia (meledina 1973, p. 51, plate 4, fig. 3) resembles some individual variants of the greenland faunas in size and coiling, but the style of the ribbing points rather at later forms from the pompeckji standard zone, which occur in the same bed (meledina 1973, p. 104, bed 6, 0.4–0.7 m). pompeckji standard zone the pompeckji standard zone is herein divided into four new subzones: the intermissus subzone, the carlsbergensis subzone, the gracilis subzone and the episcopalis subzone (fig. 30). intermissus subzone the intermissus subzone comprises seven faunal horizons: po-1–po-7, forming the group of cranocephalites intermissus sp. nov. this is a succession of transients sharing salient morphological features that mark a significant change from preceding taxa. the macroconchs are larger, more inflated, have depressed whorl-sections, and are strongly ribbed to the end. they have homeomorphs at the levels of po-11 – po-13 and again at po-19 – p-20, but are stratigraphically clearly separated. po-1: cranocephalites carolae sp. nov. sensu stricto, trans α plate 6, figs 1–4; plate 8, fig. 3 (all [m]); fig. 34. types. holotype jhc 5205 (plate 6, fig. 1), paratypes jhc 5215, jhc 5229, jhc 5234 (plate 6, figs 2–4), all part of the reference collection, section c4, bed 12b (fig. 21). material and distribution. this is another example of an abundant association of adult macroconchs in a beautiful characteristically dark-coloured phosphatic preservation, localised at one thin horizon clearly recognisable over an area of not more than about a square kilometre (sections c2–c4). the reference collection, section c4 (fig. 21): 31 [m]. section c3, bed 3b (fig. 20): 14 [m]. section c2: 1 [m]. total: 46 (all [m], microconchs not recognised). etymology. named to acknowledge with gratitude the assistance over 30 years, both in the field and in the collections, of carol st john payne. description. macroconchs of medium size, section rounded and moderately depressed, umbilici wider and more open than in earlier species. phragmocone densely, sharply and dmax dph h w u jhc 6023: jhc 6024: dmax, dph in mm 0.6555 0.7265 90 100 dmax dph h w u jhc 6031 [m]: jhc 6047 [m]: dmax, dph in mm 0.57 0.1236 0.5158 (est.) 29 61 finely ribbed (plate 8, fig. 3) in the style that then remains essentially unchanged over the whole range of successive transients in the pompeckji standard zone. adult body chamber only mildly variocostate, the primary ribbing persisting to the end, the secondaries passing over the venter with only mild attenuation in strength (plate 6, figs 2, 4), if any; the finer-ribbed variants (plate 6, fig. 3) very regularly ribbed, the more coarsely-ribbed ones tending to become somewhat bullate (plate 6, fig. 4). the variability may be seen in the reference collection photographed in the field and shown in fig. 25. microconchs still unknown. dimensions. po-2: cranocephalites carolae sp. nov. trans β plate 7, figs 1–9 (all [m]) primary representative specimen of transient β. jhc 6430, section d3 (plate 7, fig. 6). secondary representative specimens of transient β. jhc 6432 (plate 7, fig. 1), jhc 5336 (plate 7, fig. 2). material and distribution. yet a further example of a highly localised (sections d2–d3), sharply-defined horizon with an abundant assemblage of finely-preserved mature macroconchs. in this case, however, it is the phragmocones that are predominantly preserved, in a light orange-coloured calcite, reflecting the concretions in which they occur. reference collection: section d2 (fig. 22), on the top of a concretionary, calcareous sandstone c. 2 m below the base of the residual wedge of the clinoform bed of the ugleelv member at katedralen: 38 [m]. section d3, 0.5 km to the west of section d2, in a layer of large, well-separated concretions, the only resistant feature in some 20 m of soft, dmax dph h w u jhc 5205 (ht): mean values, others: <dmax> 81 ± 1.8 σ = 9.0 (11%) n = 26 <dph> 55 ± 1.2 σ = 7.0 (13%) n = 26 <wph> 0.59 ± 0.00 σ = 0.06 (10%) n = 26 dmax, dph in mm 0.485483 0.59 0.20 fig 34. a field collection made on a wide exposure of horizon po-1, cranocephalites carolae sp. nov. hammers for scale. 62 recessive mudstones, 11 m below the clinoform bed: 17 [m]. total: 55 (all [m], microconchs not recognised). description. macroconchs share the relatively open umbilicus of trans α but are somewhat more compressed, particularly in the inner whorls; adult body chambers more strongly variocostate, the secondaries more strongly differentiated (plate 7, figs 1, 2, 6). dimensions. po-3: cranocephalites intermissus sp. nov. trans α plate 8, figs 4, 5 primary representative specimen of transient α. jhc 6510 (plate 8, fig. 5). secondary representative specimen of transient α. jhc 6512 (plate 8, fig. 5). material and distribution. another highly localised horizon, producing beautifully preserved specimens few of which retain anything of the body chamber. reference collection: teebjerg, slopes above falkeelv, section c3, bed 5, 4 m above po-1 (fig. 20): 20 [m]. statuebjerg, section b4, bed 8: 10 [m]. total: 30 (all [m], microconchs not recognised). description. larger, more round-whorled and inflated than taxa in po-1 and po-2, narrower umbilicus; more densely and finely ribbed to greater diameters. the ribbing coarsens and modifies somewhat on the adult body chamber, the primaries becoming more strongly differentiated. dimensions. po-4: cranocephalites intermissus sp. nov. sensu stricto, trans β plate 9, figs 1–4 [m], fig. 5 [m] 1984 cranocephalites intermissus heinberg & birkelund, p. 376, fig. 13. 1993 cranocephalites sp. nov. a [intermissus ms], callomon, p. 96 [fauna 3]. holotype [m]. ggu 139115a, birkelund and heinberg collection 1974, katedralen, section d9, bed 15. allotype [m]. jhc 5119, section d8, bed 17d (fig. 25), clinoform bed of the ugleelv member at katedralen, highest part. material and distribution. the name was introduced at a time when this was the only fauna, then newly discovered, lying between those of the c. indistinctus group below and those of the c. pompeckji/gracilis/maculatus horizons above. the type, the reference and largest collection came from a more recessive, mudstone interbed in the middle of the clinoform bed on the west side of the katedralen canyon, section d9, 9 m above its base and 16 m below its top: 14 [m]. found also on the opposite side of the canyon, to the east, section d8 (fig. 25), similar stratigraphic level: 1 [m], 1 [m]; on the opposite side, to the north, of ugleelv, section e1 (fig. 27): 14 [m]. there are further scattered records, all within the clinoform bed, serving to confirm that the time needed to build this unit was relatively short, within the typical time interval between successive faunal horizons. the state of preservation is generally rather poor. total: 41 (c. 40 [m] + 1 [m]). description. typically depressed, round-whorled macroconchs, with fine, dense ribbing, the primaries rising on wellrounded umbilical shoulders; sculpture of the adult body dphdmax h w u jhc 6510 (wholly septate but phragmocone complete; plate 8, fig. 5): mean values, reference collection: <dph> 65 ± 0.8 σ = 3.7 (6%) n = 21 <wph> 0.59 ± 0.02 σ = 0.050 (8.5%) n = 21 dph in mm 0.4672 0.56 0.14 0.5069jhc 6512 (plate 8, fig. 4): 0.58 0.11 dmax dph h w u jhc 6430: mean values, reference collection: <dmax> (94) (estimated from <dph> x ρ(π)) <dph> 62.3 ± 0.9 σ = 4.0 (6.4%) n = 22 <wph> 0.54 ± 0.01 σ = 0.053 (10%) n = 39 ρ(π) 1.34 σ = n = 6 *incomplete body chamber dmax, dph in mm 0.486581* 0.57 0.16 63 chambers only moderately modified, although the variability of all characters is considerable. dimensions. po-5: cranocephalites sp. (u-5) (not figured) this is the lowest of the biospecific transients of crano cephalites that has been identified in the highly condensed succession of ironstone that lie above the clinoform bed on the ne slopes of katedralen, along the sw side of ugleelv above its junction with ræveelv, sections d2–d10 and sections e1–e3. the development of these lenticles is shown on fig. 35. material. the fauna has been found at only a single locality, section d8 (fig. 25) in the katedralen canyon and the recoverable assemblage of moderately well-preserved material consists of only three macroconchs. these are somewhat phosphatised and crushed, only the body chambers being preserved. the bed was fairly fossiliferous and fragments of ammonites quite abundant. the forms are, however, so different from those below and above, and their positions immediately on top of the clinoform bed so secure, that their differentiation into a separate faunal horizon seems justified. nevertheless, its position in the succession must remain somewhat tentative. description. a relatively small form, the more evolute variants tending to a depressed but subquadrate whorl-section, the more involute forms inflated and depressed as in c. intermissus. strongly ribbed to the end, with some mid-ventral weakening of the secondary ribbing on the body chamber. dimensions. po-6:cranocephalites sp. aff. intermissus (u-6) (not figured) material. another assemblage from a well-characterised stratigraphical level at a single locality, section e3, bed 2, in an ironstone band in soft mudstones 0.5 m above the diplocraterion habichi burrowed top of the clinoform bed of the ugleelv member (fig. 35b). the material is plentiful but consists mainly of phragmocones. total: 19 (all [m]). description. similar to c. intermissus but markedly more evolute and compressed in the inner whorls, foreshadowing the later forms of the c. pompeckji group, its main component, po-8, occurring in the same section only 0.5 m higher. the variability has also increased. dimensions. po-7:cranocephalites sp. aff. intermissus (u-7) (fig. 36) material. another assemblage from a single locality, section d6, in the lowest 0.4 m of non-ferruginous sandy mudstones or soft fine-grained sandstone resting directly on the clinoform bed of the ugleelv member at katedralen. these mudstones have a total thickness of 1.9 m. no other ammonites were found in the higher part of the bed. total: 17 (all [m]). description. the collection consists of well-preserved moulds of the adult body chamber, complete with peristome, providing an unusually good insight into the mature stages of dmax [m] dph h w u ggu 139115a (ht; incomplete, 0.6 whorl body chamber preserved; plate 9, fig. 1): mean values, reference collection: <dph> 57 ± 2 σ = 7 (12%) n = 12 <wph> 0.66 ± 0.02 σ = 0.06 (9%) n = 13 dmax [m] dph h w u dmax, dph in mm jhc 5119 (at; 0.75 whorl body chamber, complete; plate 9, fig. 5): at d = 25: d, dmax, dph in mm 0.425568 0.65 0.17 0.51 2035 0.58 0.16 dmax dph h w u jhc 5125 (a complete adult body chamber, evolute variant with peristomal constriction): at d = 55: d, dmax, dph in mm 50 ~0.55 75 dmax dph h w u jhc 5447 (inflated): jhc 5444 (compressed): mean values: <dph> 57.2 σ = 8.2 (14%) n = 12 <wph> 0.58 σ = 0.082 (14%) n = 12 dph in mm 0.4864 0.70 0.14 0.4758 0.50 0.15 64 ontogeny. the phragmocone tends to be lost by crushing in the compressed variants but may be retained in the inflated ones. the assemblage is marked by a high variability in the inflation of the shell, with more compressed variants now dominating. it illustrates yet another example of the pattern of phyletic change of morphology widely observed in ammonites, here the change from inflated to compressed whorl-section of the shell by proterogenesis, the appearance of new characters in the young, subsequently spreading to the later stages of ontogeny (schindewolf 1925). the assemblage is marked by high variability in the inflation of the shell, with more compressed variants now dominating (fig. 36a, b). the average diameter of the adults continues, however, to be closely constrained. the ribbing persists strongly to the end, the secondaries crossing the venter on the adult body chamber in undiminished strength in inflated variants, or fading to smoothness midventrally on the early part of the body chamber. this range of variability generates variants that are close homeomorphs of almost all the other transients of cranocephalites, including c. intermissus, maculatus, pompeckji, and even indistinctus as an extreme. conversely, no single specimen characterises the assemblage as a whole. it seems to epitomise the genetic, phenotypic plasticity in expressions of the genome of ammonites as a whole. yet the stratigraphic ‘purity’ of the biohorizon seems in no doubt; this is not a heterogeneously condensed assemblage. dimensions. 40 38 36 34 30 28 20 16 14 12 10 1-8 18 500 450 35c 35b 400 arcticoceras jhc 6338–6340 arcticoceras cf. ishmae arcticoceras sp. arctocephalites greenlandicus, arctocephalites cf. freboldi (j-12) arctocephalites arcticus (j-9) jhc 6331–6337 jhc 6236–6330cranocephalites spp. (see fig. 35b) ræveelv: section e3 alithostratigraphy ammonite sample no. m a . s l. be d no . u gle el v m em be r fb -3 fb -4 fb -2 u g4 fb -1 pe lio n fm fo ss ilb je rg et f or m at io n h ar ee lv fo rm at io n fig 35. section e3 at ræveelv, the origin of the reference collection of c. carlsbergensis trans α (po-10), and c. transitorius trans α (po-17), shown in a generalised log (a) and in a detailed log (b: lower part. c: upper part), see facing page and following page. for location, see fig. 9; for legend, see fig. 12. dmax dph h w u jhc 6351 (cf. c. intermissus): jhc 6348 (cf. c. pompeckji): mean values, reference collection: <dmax> 81.0 σ = n = 12 n = 12<dph> 56.0 0.50 σ = 4.2 (7.6%) σ = 0.091 (18%) n = 12 (0.42 < w < 0.67)<wph> dmax, dph in mm 0.476480 0.67 0.14 0.495581 0.47 0.17 65 400 390 2 8 6 7 b a 5 34 9 20 18 16 14 12 10 410 420 430 440 arctocephalites arcticus (j-9) jhc 6331–6337 cranocephalites transitorius β (po-18) cranocephalites transitorius α (po-17) (not collected)arctocephalites greenlandicus arctocephalites cf. freboldi (j-12) cranocephalites gracilis (po-15) cranocephalites carlsbergensis α (po-10) cranocephalites pompeckji (po-8) cranocephalites sp. (u-6) (po-6) jhc 6298–6330 jhc 6289–6297 jhc 6270–6288 jhc 6258–6269 jhc 6248–6257 jhc 6236–6247 ræveelv: section e3 (lower part) blithostratigraphy ammonite sample no. m a . s .l be d no . u gle el v m em be r u g3 u g2 u g4 fb -1 fb -2 fo ss ilb je rg et f or m at io n pe lio n fo rm at io n 66 440 40 38 36 34 32 30 28 26 24 22 20 450 460 470 480 arcticoceras arcticoceras sp. (not collected) arctocephalites sp. (not collected) arctocephalites greenlandicus, arctocephalites cf. freboldi (j-12) (not collected) arcticoceras cf. ishmae (not collected) jhc 6338–6340 ræveelv: section e3 (upper part) clithostratigraphy ammonite sample no. m a . s l. be d no . fb -2 fb -3 fb -4 fo ss ilb je rg et f or m at io n h ar ee lv fm fig 35. for caption see previous pages. 67 carlsbergensis subzone the carlsbergensis subzone of the pompeckji standard zone includes seven faunal horizons, po-8 – po-14 (fig. 30), described below under two groups, po-8 – po-9 and po-10 – po-14. po-8 – po-9: the group of cranocephalites pompeckji madsen – furcatus spath po-8:cranocephalites pompeckji (madsen 1904) plate 10, figs 1–6; plate 11, figs 1, 2 [m]; plate 11, fig. 3 [m] 1904 macrocephalites pompeckji madsen, p. 189, plate 8, figs 5, 6 (lt). ?1913 macrocephalites pompeckji madsen – sokolov, p. 61, plate 1, figs 1, 2 (novaya zemlya). 1929 macrocephalites pompeckjimadsen – rosenkrantz, p. 146. 1930 macrocephalites pompeckji madsen – frebold, p. 111. 1932 cranocephalites pompeckji (madsen) – spath, p.16, non vars. laevis (plate 5, fig. 3, = arctocephalites sp.), rustica (plate 4, figs 9, 10; plate 13, fig. 1, = c. furcatus), intermedia a b c 1 cm d ba fig. 36. cranocephalites sp. aff. intermissus (u-7), section d6, bed 1a, faunal horizon po-7. a, b: side view and ventral view of jhc 6348, a typical slender variant. c, d: side view and apertural view of jhc 6351, the most inflated variant. all natural size. 68 (plate 5 fig. 7, = c. furcatus), costata (plate 5, fig. 6; plate 13, fig. 1, = c. furcatus). 1932 cranocephalites vulgaris spath, p. 20, plate 1, figs 2, 3 (ht), 6 (suture only). cranocephalites vulgaris spath var. robusta spath, p. 22, plate 2, fig. 1. cranocephalites vulgaris spath var. densicostata spath, p. 21, plate 4, fig. 1. cranocephalites vulgaris spath var. compressa spath, p. 21, plate 5, fig. 1. cranocephalites aff. vulgaris spath, plate 2, fig. 4. cf 1962 cranocephalites pompeckji (madsen) – voronets, p. 33, plate 7, fig. 1a, b; plate 23, fig. 3. cf 1962 cranocephalites vulgaris spath – voronets, p. 32, plate 22, fig. 2a, b, (northern siberia, anabar). ?1973 cranocephalites pompeckji (madsen) – meledina, p. 43; non plate 2, figs 3a, b, 4a, b; plate 3, fig. 3a, b (northern siberia, anabar). cf 1973 cranocephalites pseudogracilis meledina, p. 48, plate 5, fig. 3a, b (ht). 1985a cranocephalites pompeckji (madsen) – callomon, p. 56, text-figs 4, 8d [m], 8d [m]. 1993 cranocephalites pompeckji (madsen) – callomon, p. 98 (fauna 7). lectotype. mguh 297; madsen (1904), plate 8, fig. 6a, b, designated spath (1932), p. 17; refigured here; plate 10, fig. 2a, b. material and distribution. the material of madsen (1904) from ‘ammonite mountain’, presumed to be the rising plateau north of sortehat (i.e. south-west slope of dusén bjerg, fig. 3; higgins 2010, see figs 3, 11 foreground): the type and three other identifiable former syntypes (madsen 1904, plate 8, fig. 5 and two others in mguh, not numbered); section 23, = section e3 (figs 3, 9, 35b) on the opposite side of ræveelv to the west, 14 [m], 2 [m]; the north slopes of katedralen, section 20 = section d5 (figs 3, 9, 24), several collections including ones made by t.m. harris (c. vulgaris in spath 1932), f. surlyk (in 1970), t. birkelund and j.h. callomon (in 1971), t. birkelund and c. heinberg (in 1974): c. 90 [m], 7 [m]. total: >100 [m] + 7 [m]. yet another species found in abundance over only a very restricted area, and it is ironic that the only four identifiable specimens brought back by nordenskjöld should have been from just this stratigraphically highly localised biozone. at section e3, perhaps only 3 km to the south-west of ‘ammonite mountain’ (fig. 9), well-preserved ammonites from other biohorizons spread over the ground greatly outnumber those of c. pompeckji. what appears to be the fauna characterising the po-8 faunal horizon has been found at only one other locality in east greenland, antarctic havn. the material is, however, poorly preserved and the identifications must be tentative. description. the intraspecific variability among the macroconchs is relatively low and its range is illustrated in the plates and in callomon (1985a, text-fig. 4). compared with the older forms, the size is reduced, the coiling is less depressed in whorl-section, the whorl-sides tending to be flatter, the umbilicus opening at an earlier stage. the ribbing is dense and subdued, both the primaries and secondaries persisting with little modification to the end of the adult body chamber but tending to weaken and, in extreme variants, to fade mid-ventrally. a microconch with nearly complete body chamber is shown in plate 10, fig. 4, and apart from a markedly flexuous, forward-projected ribbing, it differs little from the nuclei of macroconchs at comparable diameters. not enough specimens of the microconchs of any species of cranocephalites have been found to map out their variabilities. dimensions. discussion. the discrepancy between the present interpretation of the species of madsen (1904) and that of spath (1932) brings out in acute form the conflicts that can arise between morphospecific and biospecific classifications, between comparisons of single specimens as opposed to internally isochronous but variable assemblages. on the one hand, spath (1932) picked out very successfully a coherent assemblage from the only material at all localised both in place and horizon at his disposal, that collected by t.m. harris from katedralen at 370 m. his sketch of the section was reproduced by spath (1932, p. 128, text-fig. 12). to this assemblage of 11 specimens, which is now unambiguously in assemblage po-8, he gave the name c. vulgaris. on the other hand, he had also to accommodate dmax dph h w u mguh 297 (lt; only 0.3 whorl body chamber preserved; plate 10, fig. 2): cf. jhc 4087 (complete adult, large variant; plate 10, fig. 1): mean values, reference collection: <dmax> 73.5 ± 1.8 σ = 11.5 (16%) n = 40 <dph> 50.1 ± 0.1 σ = 5.0 (10%) n = 45 <wph> 0.54 ± 0.01 σ = 0.043 (8%) n = 46 ρ(π) 1.35 dmax, dph in mm 0.506069 0.48 0.13 0.505885 0.50 0.13 69 the only two or three identifiable specimens in the old collections, essentially without provenance, that also carried the oldest name, the syntypes of c. pompeckji (madsen 1904). in this, he saw a closer resemblance of these syntypes to a number of other poorly-preserved specimens collected by a. rosenkrantz at hjørnefjeldet located considerably farther north (figs 1b, 4) than to his c. vulgaris. with the stratigraphic information now available, the matter can be clearly resolved. the assemblage from hjørnefjeldet is quite distinct from that of c. vulgaris at katedralen. it belongs to fauna po-9, that of c. furcatus spath 1932, discussed further below. regarded as an isolated variant, the type of c. pompeckji could probably be fitted equally well into either c. vulgaris (po-8) or c. furcatus (po-9), although spath himself admitted that there was not a single specimen in the new collections before him that showed perfect agreement with the type (spath 1932, p. 17). the specimen from katedralen shown here in plate 10, fig. 1, fits the type of c. pompeckji almost exactly. furthermore, the assemblage of po-8 has not been found again anywhere around hjørnefjeldet and that of po-9 is missing at katedralen. and whereas the type-locality of c. pompeckji, ‘ammonite mountain’, is at most 5 km from katedralen, it is 35 km from hjørnefjeldet. the biospecific assemblages po-8 and po-9 are both well characterised, but which should carry the morphospecific name c. pompeckji? both morphological and stratigraphical evidence point now overwhelmingly to po-8, the type assemblage also of c. vulgaris – which makes this junior synonym of c. pompeckji. some of the specimens referred to c. pompeckji or c. vulgaris by meledina (1973) agree with the types in size and coiling but not in the style of ribbing, which is consistently too coarse and strong, especially on the venter. some 34 specimens of cranocephalites were figured by meledina (1973), all but two of which came from two areas in the region of the anabar river in northern siberia (ury ung –tumus and anabar bay) about 50 km apart. the successions are summarised in two sections (meledina 1973, p. 102–105) which overall are very similar (meledina 1973, fig. 35). the specimens appear to have come mostly from a layer of concretions in a bed of clay or silt (bed 6) that is 16 and 10 m thick at the two localities, respectively. their stratigraphic positions are given as cranocephalites vulgaris zone. the 34 specimens are assigned to 14 nominal species, of which four are new and eight are based on the types of species from east greenland described by spath (1932). from what is now known for their type horizons, these greenland species range in age over a time span of 20 biohorizons, from in-6 (c. indistinctus) to po-19 (c. maculatus). it would be tempting, therefore, to conclude that the cranocephalites bed of spath (1932) is condensed, containing a mixture of fossils of many ages. but with again the experience in greenland of the variability encountered in the assemblages from single biohorizons, that of isochronous biospecies, it is equally possible to suggest a contrary interpretation. of the 34 specimens illustrated by meledina (1973), 26 would fit comfortably into the variability-distribution expected in but a single biospecies, say c. sp. a. all these specimens share a characteristic style of ribbing, no matter how large or inflated. (the exceptions are mele dina’s (1973) plate 2, fig. 2, plate 3, fig. 2: “c. furcatus”, say sp. b; plate 4, fig. 1 from the lena river, “c. nordvikensis”; plate 5, fig. 2, “c. aff. costidensus”; plate 7, fig. 4; plate 8, figs 2, 4 from the lena river, “c. gracilis”, say sp. c). taken, then, as a single biospecies, the 26 specimens of species a do not together match precisely any of the biospecific transients from jameson land, neither in the macroconchs nor in the (single) microconch (meledina 1973, plate 2, fig. 4). the closest resemblance is to the stout, coarsely-ribbed forms of c. carlsbergensis (po-11), or c. tvaerdalensis (po-13). the siberian assemblage contains variants showing an extreme degree of bullate variocostation on the adult body chamber (meledina 1973, plate 4, fig. 2, “c. laevis” and plate 9, fig. 2a, b, “c. maculatus”) unknown in greenland. it would be safest, therefore, to regard this siberian biospecies a as a distinct transient fitting into the succession somewhere in the upper carlsbergensis subzone of the pompeckji standard zone. to name it, there is an available choice of four new species of meledina (1973), of which c. spathimeledina 1973, the type species of pachycephalitesmeledina would be the most typical. another distinct biospecies of cranocephalites from northern siberia has been described by voronets (1962) as c. nordvikensis (p. 39; lt plate 6, fig. 1a, b, designated by meledina 1973). it probably includes other specimens described under a variety of names in plate 4, figs 2a, b, 3a, b; plate 6, fig. 2; plate 9, fig. 1; plate 12, fig. 3 and plate 13, figs 2a, b, 3a,b (voronets 1962). taken together, they have no close match in east greenland. po-9:cranocephalites furcatus spath 1932 plate 11, figs 4, 5; plate 12, figs 1–4 [m]; plate 11, fig. 6; plate 12, fig. 6 [m] 1932 cranocephalites furcatus spath, p. 30, plate 6, fig. 1 (ht). cranocephalites furcatus spath var. pygmaeus spath, p. 30, plate 6, fig. 2 [m?]. 70 cranocephalites pompeckji (madsen) – spath, p. 16, plate 4 fig. 10, plate 13, fig. 1[m]. cranocephalites cf. pompeckji (madsen) – spath, p. 19, plate 3, fig. 3 [?m]. cranocephalites pompeckji (madsen) var. costata – spath, p. 18, plate 5, fig. 6. cranocephalites pompeckji (madsen) var. intermedia – spath, p. 18, plate 5, fig. 7a, b. cranocephalites pompeckji (madsen) var. rustica – spath, p. 18, plate 4, fig. 9a, b. cf 1953 arctocephalites (cranocephalites) maculatus spath var. rusticus nov., donovan, p. 83, plate 16, figs 1–3. 1993 cranocephalites furcatus spath – callomon, p. 98. non 1973 cranocephalites furcatus spath: meledina, plate 2, fig. 2; plate 3, fig. 2. holotype. mguh 9191, refigured here (plate 11, fig. 4). type locality and distribution. the type collection and the other specimens described by spath (1932, see synonymy) were spot collections brought back by a. rosenkrantz from his ‘mt. hjørnefjeld’, ‘locality 2b, 740 m’. a sketch of the section is shown in rosenkrantz (1934, p. 92, fig. 45) and in spath (1932, p. 133, text-fig. 14). attempts to find this precise locality again during mapping in 1971 were unsuccessful. the highest point in the area does not rise above 660 m. but the po-9 beds are widely exposed in the region, particularly around centralbjerg to the west of hjørnefjeldet (fig. 4, section 35) and thence further north, and the two levels 20 m apart with ammonites recognised by rosenkrantz were readily found again and new collections made. the lower is the type-horizon of c. furcatus, the higher that of arctocephalites arcticus (newton 1897) and a. nudus spath 1932, fauna 9 in callomon (1993, p. 99). at the time the callomon (1993) review was written, the relative positions of c. furcatus and c. pompeckji were not definitely known, for whereas c. furcatus had at that time not been found in the region of katedralen and ugleelv, c. pompeckji was unknown in the region around hjørne fjeldet and centralbjerg. after much searching, c. furcatus, now horizon po-9, was found in 1994 on the eastern slopes of katedralen, in section d2, bed 10 (fig. 22), at a level of a thin ironstone bed immediately above the clinoform bed, horizon po-2 of the ugleelv member, and directly overlain by another ironstone with the fauna of po-10. there was no sign of the horizons po-3 – po-8, so the succession at section d2 is here highly incomplete, but po-9 must lie above po-8. material. centralbjerg region: c. 60 [m], 9 [m]; katedralen, 18 (all [m]). description. the photographs of the holotype belie its description. the specimen essentially shows only one side of a macroconch body chamber. it shows, however, the characteristic features of the species: the regular, strong but blunt well-spaced ribbing, the primaries dividing mostly into pairs of secondaries of comparable strength. a specimen that closely matches the type is shown for comparison in plate 11, fig. 5, with part of the body chamber removed to reveal the similarly coarse but sharp ribbing on the inner whorls. these characters are also evident in the four specimens shown on plate 12, figs 1–4. of these, plate 12, fig. 2 illustrates also the rather flat-sided, subquadrate whorl-section of typical variants and the loss of ribbing on the venter of the adult body chamber. the range of inflation is reflected in plate 12, fig. 1, which illustrates the extreme. one of the smallest variants, diameter 65 mm, was figured as var. pygmaeus by spath (1932); the maximum size can reach 95 mm. overall, the size and coiling resemble those of the immediately underlying c. pompeckji as discussed above, but the style of ribbing is quite distinct. a microconch is shown in plate 11, fig. 6a, b; the specimen figured by spath (1932, plate 3, fig. 3) is probably another microconch. dimensions. po-10 – po-14: the group of cranocephalites carlsbergensis callomon the lower three faunal horizons of this group (p-10 – po12) are defined by three transients (α , β, γ) of the taxon (cranocephalites carlsbergensis), the type of which is represented by transient β. po-10 – po-12: cranocephalites carlsbergensis callomon 1975 cranocephalites carlsbergensis (sp. nov. ms) callomon, p. 383, fig. 6a, b; 6c [m]. 1985a cranocephalites sp. nov. a, callomon, text-fig. 8c [m], 8c [m]. dmax dph h w u mguh 9191 (ht): cf jhc4172: mean values, reference collection: <dmax> 75 ± 1.5 σ = 6.0 (8%) n = 17 <dph> 48 ± 1.3 σ = 5.4 (11%) n = 15 dmax, dph in mm ~0.49~5986 ~0.15 0.49 0.565074 0.14 71 1993 cranocephalites sp. nov. b [carlsbergensis ms], callomon, p. 97 [fauna 5]. holotype. jhc 1751 (in callomon 1975, fig. 6c), by monotypy. the type comes from the faunal horizon of transient β (po-11), see below. these three faunal horizons contain very similar forms but are found so far only in three non-overlapping areas. their relative succession is therefore surmised indirectly on the basis of lateral lithostratigraphical correlations rather than on direct sequential observation in a single section. po-12 (highest): ugleelv, east of katedralen, around statuebjerg, sections b1–b3 (figs 9, 18, 19): c. carlsbergensis transient γ. po-11: trefjord bjerg, sections 39a and 39b (figs 4, 29): c. carlsbergensis transient β. po-10 (lowest): katedralen and ræveelv, sections d2–d4, e3 (figs 9, 35b): c. carlsbergensis transient α. the position of po-12 above po-10 seems assured, but the position of po-11 relative to these is still conjectural. it is placed here in the middle for convenience. up to 1994, even the position relative to c. pompeckji, po-8, was uncertain, and po-8 was then still placed above po-10 – po-12. the three transients are so similar that only large collections allow distinction. it would suffice for most problems of correlation with material from other localities, therefore, simply to refer to the three transients collectively as c. carlsbergensis. the type horizon of the species sensu stricto is po11 (trans β). po-10: cranocephalites carlsbergensis trans α figs 37–38 representative specimens of trans α. jhc 6263, [m], from the reference collection, section e3; jhc 5411, [m], from section d2, bed 11 (fig. 22). material and distribution. ugleelv (fig. 9), section d2, bed 11: 20 [m], 1 [m] (figs 22, 37). further ugleelv setions: d3: 4 [m]; d4 (fig. 23): 13 [m]; e3, reference locality and horizon (fig. 3, section 23; figs 9, 37), bed 7: 23 [m]. total: 61 (60 [m] + 1 [m]). the exposures in sections d2 – d4 were somewhat scree-covered, so that it was not always easy to collect the material precisely in situ. the ammonites of po-10 were abundant, however, and could be easily recognised again even in sections in which the thin ironstones in which they occurred were barely recognisable. in section e3, the bed crops out on a gently rising slope and the ammonites were lying in a narrow strip well separated from those of the fauna below (po-8) and above (po-15). the preservation has conserved mostly the mature body chambers, with peristome; the phragmocones tend to be crushed. description. slightly smaller than trans β and more strongly and persistently ribbed, the venter more highly arched, but the styles of ribbing and its variability match those of trans β in almost all respects. the differences lie at the extremes: no single specimen epitomises the differences between trans α and trans β. the microconchs continue to be extremely rare, as are specimens that were demonstrably juvenile. dimensions. po-11: cranocephalites carlsbergensis sensu stricto, trans β plate 12, fig. 5; plate 13, figs 1–4 [m], 5 [m]; fig. 39 types. holotype, jhc 1751 (in callomon 1975, see above); paratype i, jhc 1754 [m], plate 13, fig. 1; paratype ii, jhc 1756; allotype [m], jhc 1768 (plate 13, fig. 5). material and distribution. almost the whole of the material on which this transient is based came from two adjacent localities on the north and north-east ridges of trefjord bjerg (figs 4, 29). the first collection was made in 1958 on the north ridge, section 39b (figs 29, 40), and is the reference collection that includes the types: 47 [m], 1 [m]. its appearance at the time of collection is shown in fig. 39. a second collection was made in 1974 by c. heinberg and t. birkelund in section 39a, above lepidopteris elv, only 1.3 km to the south-east of 39b: 15 [m]. scattered finds were made elsewhere in the region during mapping. description. a large species, the inner whorls involute with rounded, depressed section and tight umbilicus (plate 12, fig. 6; plate 13, figs 3, 4), much as in the other species of cranocephalites. the adult body chambers contract with marked uncoiling of the umbilical seam; ribbing moderately variocostate, characterised by coarse, blunt primary ribbing tending to circumbilical bullae and secondary ribdmax dph h w u jhc 6263 (fig. 37a, b): mean values, reference collection: <dmax> 83.9 ± 1.9 σ = 7.7 (9.2%) n = 18 n = 20<dph> 59.3 ± 1.2 0.55 ± 0.02 σ = 5.4 (9.1%) σ = 0.084 (15%) n = 17<wph> dmax, dph in mm 0.486289 0.55 72 ba 1 cm c fig. 37. cranocephalites carlsbergensis trans α (po-10), representative specimens. a, b: side view and ventral view of the primary representative specimen, jhc 6263 [m], ræveelv, section e3. c: secondary representative specimen, jhc 5411 [m], katedralen e, section d2. all natural size. fig. 38. a field collection of cranocephalites carlsbergensis trans α (po-10) from ræveelv, section e3, bed 7 (fig. 35b). hammers for scale. 73 10 cm trefjord bjerg sf of section 39b pf fig. 39. part of the type collection of cranocephalites carlsbergensis sensu stricto trans β (po-11), as collected in section 39b, bed 11 (fig. 29), on the north ridge of trefjord bjerg (fig. 40) on display on the snow. fig. 40. the northern slopes of trefjord bjerg, showing section 39b (fig. 29), the sunlit ridge rising from the col at the left to the summit. the resistant formation forming the saddle of the col and running to the right under the snow-field is the ostreaelv formation (of) of the neill klinter group (toarcian). above it, the recessive sortehat formation (sf, c. 105 m thick), largely covered by snow running up to the base of the first massive bluff which marks the lower part of the pelion formation (pf). the type-horizon of cranocephalites carlsbergensis (po-11) lies about half-way between this bluff and the summit (white arrow). the highest sandstones, forming the plateau, are probably in the ishmae zone, horizons j15 – j16 (figs 8, 30). 74 bing that fades, leaving the ventral side of the body chamber smooth. quite close homeomorphy with po-1, c. carolae in the coarser-ribbed variants, and with po-19, c. maculatus, in the finer-ribbed forms. dimensions. discussion. some other nominal species created by spath (1932) bear some resemblance to c. carlsbergensis. three of them (c. inversus, c. inconstans and c. subbullatus) are based on material from antarctic havn that is so poorly preserved that the names are best set aside as nomina dubia. po-12: cranocephalites carlsbergensis trans γ plate 16, fig. 6 representative specimens of transient γ. ggu 137939.1, from statuebjerg, section 13/b1, bed 27; ggu 137973.1, teebjerg, loose. material and distribution. this fauna characterises a sharplydefined thin concretionary bed of soft sandstone lying at the top of some 5 m of silty mudstones. the weathered out ammonites were lying in profusion on the gentle slopes but their preservation in a characteristic hard, grey-weathering, somewhat phosphatic very fine-grained rock reveals their source. in most cases only the body chambers have survived more or less undistorted, the phragmocones having been crushed. it was difficult to find a single specimen with both uncrushed body chamber and phragmocone to act as type. uncrushed phragmocones do occur but then lack the body chambers (plate 16, fig. 6). such inner whorls are however so similar to those of other species that they contribute little to the characterisation of the assemblage. section b1: c. 100 [m]; section c1: 16 [m] + 1 [m]. the bed makes an excellent marker in the area of statuebjerg, sections b1 – b3 (figs 9, 18, 19). it can just be recognised on teebjerg, section c1, 2 km to the northwest, but no further. together with the whole of the indistinctus standard zone and the rest of the pompeckji standard zone, it has been totally lost at taubjerg, only 3 km to the south-east (fig. 9). at statuebjerg, it marks the top of the unit fb-1 of the fossilbjerget member and is there followed immediately by the mudstones of fb-2, probably greenlandicus standard zone horizon j-11. description. somewhat smaller than trans β, (c. carlsbergensis sensu stricto), body chambers more evolute, more densely and finely ribbed, fading earlier, leaving the ventral side more commonly smooth. some variants are homeo morphs of c. gracilis in po-15. dimensions. the variability distributions have some unusual tails. on the one hand, there are a few variants at the top end of the scale that have very large, inflated phragmocones, dph c. 70 mm, w = 0.61–0.68. at the other end, there are small forms, dmax c. 55–60 mm, also mature adults, in which the length of the body chamber is reduced from its otherwise ubiquitous 0.75 of a whorl to 0.55–0.60 of a whorl. po-13 – po-14 faunal horizons po-13: cranocephalites tvaerdalensis alsen 2015 plate 14, figs 1–4 [m], fig. 41 [m] cf 1962 cranocephalites (?) nordvikensis (pars), voronets, plate 4, fig. 3a, b. 2015 cranocephalites tvaerdalensis alsen, plates 1–5 (appendix 1). types. holotype mguh 31377 (ex ggu 522023a); allotype mguh 31378 (ex ggu 444903). type material from tværdal, geographical society ø (appendix 1). material and distribution. known in jameson land only from two closely adjacent sections at katedralen, section d4 (bed 12b; fig. 23) and section d5 (bed 5c; fig. 24), dmax dph h w u ggu 137939.1: at d = 60: ggu 137973.1: mean values, reference collection: *inner whorls crushed d, dmax, dph in mm <dmax> 91.2 ± 2.9 σ = 9.3 (14%) n = 11 <dph> 66.6 ± 1.2 σ = 8.4 (13%) n = 12 <wph> 0.55 (±0.02) σ = 0.083 (15%) n = 15 ~55*81 0.52 0.505583 0.50 0.18 dmax dph h w u jhc 1754 (plate 13, fig. 1): jhc 1731 (plate 13, fig. 3): mean values, reference collection: <dmax> 88.5 ± 1.4 σ = 8.1 (9.2%) n = 35 <dph> 62.7 ± 0.80 σ = 5.7 (9.0%) n = 53 <wph> 0.59 ± 0.009 σ = 0.064 (11%) n = 52 n = 2ρ(π) 1.33 (± 0.015) dmax, dph in mm 0.526092 0.60 0.13 0.525679 0.70 0.10 75 restricted in the latter to a single bed of ironstone 0.5 m thick, separated only by partings from similar ironstones of po-10 below and po-15 above. in parts of the bed the ammonites are clustered together. mature macroconchs dominate as usual, but in one case a macroconch has attached to it a cluster of two microconchs, with imprints of more (fig. 41). section d4 (fig. 23): 29 [m]; section d5 (fig. 24): 22 [m], 2 [m]. total: 53 (51 [m] + 2 [m]). description. this is probably the most strongly-sculptured, strongly-ribbed transient of all the cranocephalites of the pompeckji standard zone, particularly on the middle whorls (plate 14, figs 2, 4), although extreme variants in other transients can be similar. the adult size is average but the whorl-section markedly inflated and depressed. dimensions. comparisons. given the size of the available collections, there can be little doubt about the close synchroneity of the assemblages from geographical society ø and jameson land, the only localities at which they have so far been found, some 200 km apart. yet at both localities the occurrences are restricted to very small areas, of the order of a few square kilometres. and at both localities, the assemblages consist exclusively of mature adults, the dimorphic ratios heavily balanced in favour of the macroconchs. in jameson land they are found in thin, highly condensed facies. on geographical society ø, they occur in a thin bed in a sandy succession. the quality of the fossils indicates that the shells had a high preservation potential. the highly disjunct distributions of the ammonites seem therefore once again to point to ecological rather than taphonomic ones as the main factors determining the observed distributions. po-14: cranocephalites sp. aff. tvaerdalensis (u-8) (not figured) representative specimens of (u-8). section d4, bed 12c (fig. 23): primary representative specimen, jhc 6117 [m]; secondary representative specimens, jhc 6116, jhc 6114. material. another small collection found in a very restricted area, in sections d4 and d3 on the north slopes of katedralen (fig. 9), but in a sharply defined stratigraphical unit, an ironstone 0.3 m thick bounded by clay partings, bed 12c, underlain by po-13 (c. tvaerdalensis) and overlain directly by po-15, c. gracilis (fig. 23). the lithologies and preservations are also distinct. the specimens of po-14 are mostly fragmentary or distorted, the phragmocones either crushed or lost through bioturbation by large burrowers. the onset of the body chamber is in many cases not clear. sections d3, d4: 8 [m], tolerably complete, and c. 12 [m] more fragmentary. description. a relatively small species, resembling in coiling and compression c. indistinctus of in-6, but retaining the coarseness, strength and style of ribbing of c. tvaerdalensis, po-13. most of the forms cluster around a fairly compressed, small mean, but there are occasional variants that are larger and more inflated. the distribution appears to be skewed, but there is not enough material to test this. the primary representative specimen (jhc 6117) has been chosen to be average in size and compression with strong, coarse ribbing; jhc 6116 is similar but with denser ribbing fading on the venter, presaging c. gracilis of po-15; jhc 6114 represents the larger inflated forms. the length of the adult body chamber is typically only 0.65 of a whorl, shorter than in earlier forms. 1 cm fig. 41. a slab of sandstone (jhc 4165) with a cluster of a fragment of a macroconch and two microconchs of cranocephalites tvaerdalensis alsen 2015 (see appendix 1). mean values, reference collection: <dmax> 82.0 ± 1.7 σ = 6.5 (8.1%) n = 15 <dph> 61.1 ± 0.7 σ = 4.9 (8.0%) n = 46 <wph> 0.58 ± 0.01 σ = 0.05 (8.5%) n = 46 dmax, dph in mm 76 dimensions. gracilis subzone the gracilis subzone of the pompeckji standard zone includes six faunal horizons, po-15 – po-20, representing the group of c. gracilis spath 1932 – ornatus (spath 1932). these faunas from the higher part of the pompeckji standard zone largely share some common features: more planulate forms dominate and the ribbing is denser, finer, less differentiated, and less bullate on the adult macroconch body chamber. spath (1932) introduced two new specific names to mark what he regarded as the most striking features among the very limited selection of specimens he had before him in t.m. harris’ collection from katedralen: c. gracilis and c. maculatus. he appended separate varietal names to some of the specimens: c. gracilis var. ornata var. rotunda c. maculatus var. transitoria var. tenuis the determinations of spath (1932) had to be based purely on morphology for he had no further stratigraphical details. with the information now available, it becomes immediately clear that some of the taxonomic distinctions reflect differences of age: that the types came from different horizons. the problem then became the same as that in the case of c. pompeckji, viz. to see whether the types could be matched with topotypes from among the successive transients now recognised and hence their names used to label these transients. this was successful in all of the taxa, although in some cases the named types were by no means the most typical of their species. in stratigraphical order: po-20 (highest) c. gracilis var. ornata po-19 c. maculatus sensu stricto po-18 c. maculatus var. transitoria and var. tenuis po-15 (lowest) c. gracilis sensu stricto and var. rotunda two of the transients, po-15 and po-19, can therefore bear the names of full existing species. but the other two, po20 and po-18, contain previously described taxa that carry only varietal names: po-20, var. ornata and po-18, var. transitoria. it would be convenient to retain these names by promoting them to specific rank. this should here be possible within the rules under the provisions of article 45.6.4 of the international code of zoological nomenclature, which allows varietal names introduced between 1930 and 1961 to be construed subsequently as either of varietal or subspecific status even if called ‘var.’, unless the author explicitly stated ‘var.’ to be strictly that, an infraspecific category not of specific or subspecific taxonomic meaning in a biological sense. although spath (1932) carefully labelled his specimens as ‘var.’ he never gave an explicit account of what he meant by ‘species’. in numerous other cases he referred specimens to be ‘transitional’ from one ‘species’ to another, or ‘intermediate’ between one or another. it seems clear that he regarded his ‘species’ as little more than descriptive artefacts, and hence the distinction between ‘sp.’ and ‘var.’ to be one of degree rather than of principle. the names ‘ornata’ and ‘transitoria’ will therefore be used here as of species-group rank with authorship and date going back to spath (1932). po-15: cranocephalites gracilis spath 1932 plate 14, fig. 5; plate 15, figs 1–5; fig. 42 1932 cranocephalites gracilis spath, p. 22, plate 2, fig. 6; plate 3, fig. 1 (ht); non var. ornata, plate 2, fig. 6a, b (= c. maculatus). holotype. mguh 9166, collected from the north slopes of katedralen by t.m. harris, refigured here on plate 15, fig. 1a, b after some cleaning. material and distribution. this is perhaps the commonest and most widespread species in east greenland. the preservation is, however, everywhere poor. only the body chambers are preserved in slightly phosphatised, fine-grained, non-ferruginous indurated shale, lying free in the soft mudstones when weathered, the phragmocones having been lost by crushing (fig. 42). ugleelv, two reference collections: (1) section d4 (fig. 23), type-locality, bed 12d: 45 [m]; (2): section e3, west of ræveelv, bed 8: 19 [m]. others sections: d2 (fig. 22), d6, d8–10: 82 [m]. total: 149 (all [m]). other spot collections from mikael bjerg, section 33 (fig. 4), trefjord bjerg, section 38 (m. engkilde 1993 collection). traill ø: dmax dph h w u jhc 6117: jhc 6116: jhc 6114: mean values, reference collection: <dmax> 76.0 ± 2.4 σ = 5.9 (8%) n = 7 <dph> 52.5 ± 1.9 σ = 4.7 (9%) n = 8 <wph> 0.55 ± 0.03 σ = 0.09 (16%) n = 8 dmax, dph in mm 0.505076 0.50 4973 0.49 5795 0.74 77 svinhufvud bjerge (p. alsen 1996, 1997 collections, casp collection); geographical society ø: tværdal (p. alsen 1998, 2011 collections). description. markedly more evolute than most other forms of cranocephalites, the ribbing more delicate, denser, tending to fade, especially on the venters of the adult body chambers. the type is perhaps close to one end of the range of variability, probably picked out by t.m. harris because of its relatively good preservation. it bears some resemblance to the younger c. ornatus (po-20, cf. plate 18, fig. 3a, b) but both the stratigraphical position and the preservation support the difference of horizons. the other end of the range of variability is seen in plate 14, fig. 5. dimensions. in view of the poor preservation, not many measurements have significance. most of the specimens are complete up to or close to the peristome, so that the maximum shell-diameter remains reliable. the length of the adult body chamber is close to 0.75 of a whorl. the crushing of the shells goes up to the onset of the body chambers, however, so that estimates of whorl-width were made at a point half a whorl back before the peristome. the mean value of the shell-diameter at the last septum could then be estimated from the mean maximum diameter and a spiral half-whorl constant ρ(π) = 1.30, the value typical of the genus as a whole. fig. 42. a field collection of cranocephalites gracilis (po-15), at section e3, bed 8 (fig. 35b), made in 1996. hammers for scale. dmax dph h w u mguh 9166 (ht): jhc 6142: jhc 6128: mean values, reference collection (1): <dmax> 79.2 ± 1.1 σ = 6.6 (8.4%) n = 35 <dph> 59.5 (est.) σ = <wmax-π > 0.33 ± 0.006 σ = 0.033 (10%) n = 35 mean values, reference collection (2): <dmax> 80.5 ± 1.5 σ = 6.4 (7.8%) <wmax-π > 0.30 ± 0.004 σ = 0.015 (5%) n = 18 n = 18 dmax, dph in mm 0.456088 0.46 5375 0.44 53 0.49 0.4875 0.49 0.25 0.15 0.13 78 78 po-16: cranocephalites sp. aff. gracilis spath 1932 (u-9) plate 16, figs 1–5 ?1932 cranocephalites maculatus var. tenuis spath, p. 24, plate 4, fig. 2. representative specimen of (u-9). secion d1, bed 6: primary representative specimen, jhc 1328; secondary representative specimen, jhc 1340. material and distribution. two assemblages. one (jhc 5235–5250) from section c4, bed 12d (fig. 21) (immediately below the first ironstone of the arctocephalites beds, fb-2): 16 [m] and the other (jhc 1320–59) from section d1, bed 6: 19 [m]. total: 35 (all [m]). description. the inner whorls involute, compressed as in other species. evolute, serpenticone last whorl, umbilicus shallow, gently rounded umbilical shoulder and wall. primary ribbing subdued, rising slowly from the umbilical shoulder to maximum strength in lower flank, coarse, persisting to the end in full strength; there are typically 9–10 primary ribs on the final body chamber; bifurcating with intercalatories into coarse, subdued secondaries, fading mid-ventrally on the body chamber; prominent broad terminal constrictions. body chamber occupies between 0.6–0.65 of the last whorl. it differs from c. maculatus var. tenuis which is more involute, ribbing that is denser and finer, primaries that are stronger on the umbilical shoulder, lower level of bifurcation and sharper secondaries. it differs from c. gracilis (po15) which has similar coiling and size, but on average is more densely and finely ribbed; there is some overlap in variability, however, so close affinity is assured. dimensions. po-17: cranocephalites transitorius (spath 1932) trans α plate 15, fig. 6 representative specimen of transient α. jhc 6294 (plate 15, fig. 6), section e3, bed 9. material and distribution. collections at section d8, bed 18 (fig. 25): 8 [m] and section e3, bed 9b, (fig. 25): 9 [m]. total: 17 (all [m]). description. the material includes relatively few and poorly preserved specimens. they show a tendency toward flattened sides giving an almost subrectangular whorl cross section. the whorl is markedly involute until onset of the final body chamber where uncoiling then becomes marked. the final body chamber occupies c. 0.65 of the whorl. the ribbing density is similar to the smaller and slender variants of c. gracilis but slightly more subdued. the ventral weakening of ribbing on final body chamber commonly results in smooth or near smooth venters, but also commonly ribbing reappears in the last quarter of the body chamber before the peristome. broad shallow peristomal constrictions may occur. dimensions. po-18: cranocephalites transitorius (spath 1932) sensu stricto, trans β plate 17, figs 1–3; fig. 43 1932 cranocephalites maculatus sp. nov. var. transitoria spath, p. 24, plate 3, fig. 6a, b. holotype. mguh 9170, the specimen figured by spath (1932), collected by t.m. harris at katedralen. refigured here in plate 17, fig. 2a, b. material and distribution. collections at section d6, bed 1b: 8 [m], section e3, bed 10 (fig. 35b): 33 [m]; and section d9, bed 19: 2 [m]. total: 43 (all [m]). dmax dph h w u jhc 1328: mean values, reference collection (1), section d1: <dmax> 78.5 ± 1.5 σ = 4.7 (6.1%) n = 10 <dph> 53.1 ± 1.3 (est.) σ = 4.1 (7.7%) n = 10 <wph > 0.44 ± 0.077 σ = 0.051 (11.7%) n = 10 mean values, reference collection (2), section c4´: <dmax> 85.1 ± 3.5 σ = 11.2 (13.2%) <dph> 59.4 ± 2.2 (est.) σ = 7.0 (11.7%) n = 10 n = 10 dmax, dph in mm 0.475378 0.49 dmax dph h w u jhc 6294: mean values, reference collection: <dmax> 82.4 ± 2.3 σ = 6.4 (7.8%) n = 8 <dph> 55.3 ± 1.5 σ = 4.3 (7.8%) n = 8 <wph> 0.42 ± 0.01 σ = 0.021 (5.9%) n = 7 dmax, dph in mm 0.435880 0.39 0.16 79 description. the large material shows a large variation in size. the smaller variants overlap with the preceeding horizon po-17, having flat sides and dense, but somewhat stronger, ribbing. the larger forms develop a somewhat low oval whorl cross section with the maximum width low on the sides, close to the umbilical shoulder. the largest variants are also the thickest ones. the body chamber occupies 0.7–0.75 of the last whorl. the ribbing is generally stronger and coarser. weakening in a narrow ventral band occurs but consistent ribbing until final peristome seems to be commonest, perhaps giving the first indications of the ribbing in early whorls of the succeeding genus arcto cephalites. dimensions. fig. 43. a field collection of ammonites of horizon po-18, cranocephalites transitorius trans β, section e3, bed 10 (fig. 35b). hammers for scale. dmax dph h w u mguh 9170 (ht): mean values, collection from section e3, bed 10: <dmax> 84.3 ± 1.6 σ = 8.3 (9.8%) n = 28 <dph> 60.0 ± 1.3 σ = 6.4 (10.1%) n = 26 <wph> 0.48 ± 0.02 σ = 0.078 (16.2%) n = 25 dmax, dph in mm 0.506087 0.48 0.20 80 po-19: cranocephalites maculatus spath 1932 plate 17, fig. 4; plate 18, figs 1, 2 1932 cranocephalites maculatus spath, p. 24, plate 1, fig. 1a, b; ? plate 2, fig. 3a, b (a not closely identifiable phragmocone); non plate 3, fig. 6a, b (= c. transitorius), non plate 4, fig. 2. holotype. mguh 9154, t.m. harris collection 1926 from katedralen (spath 1932, plate 1, fig. 1a, b; refigured here in plate 18, fig. 1a, b). material and distribution. collections from section d1, bed 7: 16 [m], section d5, bed 5 (fig. 24): 24 [m] and 2 [m]. further 1 [m] from t. birkelund and j.h. callomon 1971 collection, 4 [m] from t. birkelund and c. heinberg 1974 collection, and 8 [m] from f. surlyk 1970 collection, loose from section d5. total: 55 (53 [m] + 2 [m]). description. compared with the previous horizon this fauna is larger and more inflated. the side is thickest on the midside and the venter is broad and rounded. the lower part of the umbilical wall is steep, but then becomes gently curved in the upper part. the inner whorls are serpenticone. the species is generally strongly ribbed. the venter commonly becomes smooth in inflated varieties, some specimens almost being wholly smooth except for bullate primaries low on the sides. others have strong persistent ribbing until final peristome. the final body chamber occupies about three-quarters of the last whorl. dimensions. po-20: cranocephalites ornatus (spath 1932) plate 18, figs 3, 4; plate 20, fig. 4 1932 cranocephalites gracilis var. ornata spath, p. 23, plate 2, fig. 6a, b (ht). holotype. mguh 9165, t.m. harris collection 1926, katedralen (spath 1932, plate 2, fig. 6a, b). material and distribution. collections from section d5 (fig. 24): 2 [m] loose, and section d8, bed 18 (fig. 25): 8 [m]. total: 10 (all [m]). description. the size variation almost similar to the previous faunal horizon (po-18), but much smaller than the overlying one (po-20). contains quite characteristic variants with narrow umbilicus, high gently sloping flanks and narrow rounded venter resulting in an arched whorl section perhaps giving an early indication of the evolution towards the disc-like arctocephalites. the fauna, however, still also contains forms with the broad, rounded venter typical of cranocephalites. dimensions. episcopalis subzone the episcopalis subzone of the pompeckji standard zone includes three faunal horizons, po-21 – po-23 (fig. 30). po-21: cranocephalites episcopalis sp. nov. sensu stricto, trans α plate 18, fig. 5; plate 19, figs 1, 2; plate 20, figs 1–3 holotype. jhc 4061 section d9, bed 22 at katedralen. material and distribution. collections from section d5 (fig. 24): 1 [m] in bed 6b and 2 [m] loose, section d7: 3 [m], section d8 (fig. 25): 4 [m], section d9, bed 22: 7 [m], section d10, bed 10: 2 [m], section 58, bed 35 in northern jameson land: 8 specimens. additionally 7 from f. surlyk 1991 collection, 3 specimens from m. engkilde 1993 collection. total: 35 (all [m]). description. very variable and characterised by very large, inflated megasphaeroceras-like forms (plate 19, fig. 1a, b), becoming relatively smooth on the body chamber, rejuvenation of the secondaries near peristome. includes coarse, large variants, retaining ribbing in full throughout. also comdmax dph h w u mguh 9165 (ht): mean values: <dmax> 88.7 ± 3.09 σ = 8.17 (9.2%) n = 7 <dph> 71.1 ± 4.04 σ = 10.7 (15.0%) n = 7 <wph> 0.52 ± 0.015 σ = 0.039 (7.5%) n = 7 dmax, dph in mm 0.455981 0.53 0.17 dmax dph h w u mguh 9154 (ht): mean values: <dmax> 89.1 ± 1.1 σ = 6.3 (7.0%) n = 33 <dph> 63.2 ± 0.9 σ = 5.0 (7.9%) n = 33 <wph> 0.56 ± 0.02 σ = 0.098 (17.5%) n = 30 dmax, dph in mm 0.51~5689 0.68 0.16 81 pressed forms commonly finely and densely ribbed. the body chamber is slightly larger than previous species and now occupies from 0.8 up to a full whorl. dimensions. stratigraphy. the morphological affinities are still within cranocephalites rather than arctocephalites: the rounded, inflated sections, dense ribbing retained, no compressed, high-whorled variant. but size and variability indicate affinities with a. arcticus. po-22: cranocephalites episcopalis sp. nov. trans β (not figured) material and distribution. only two specimens, fragmentary, jhc 5158, 5159, both [m], from section d8, bed 18 (fig. 25). description. resembles closely po-21 with persisting ribbing on adult body chamber with dense ribs that do not fade on the venter. it differs in being smaller and slightly less inflated. dimensions. po-23: cranocephalites sp. (u-10) (not figured) material and distribution. collections from section c3 (fig. 20), bed 7: 11 [m], section c4 (fig. 21), bed 13: 3 [m]. section d2 (fig. 22), bed 14: 6 [m] from j.h. callomon 1994 collection and 8 [m] from his 1996 collection. total: 30 (all [m]). all specimens are crushed. description. involute even on body chamber, cross-section difficult to describe due to the preservation but probably originally rounded, mostly high oval, venter mainly rounded, subquadratic in some specimens. ribbing is strong until late stages, dense, no bullates, and the ribs fade only slightly or more commonly not at all on the venter. dimensions. stratigraphy. in the sections where po-23 is recorded, the youngest underlying ammonites belong to po-15, c. gracilis (fig. 22). the ribbing with strong ribs on the body chamber and commonly no weakening on the venter indicate it to be a very late cranocephalites. c. gracilis is represented in section d8, where it is overlain by po-17 and the faunal horizons po-20 – po-22, in rapid succession. po-23 is somewhat smaller and less inflated than po-22, the latter being transient from the underlying larger and more inflated c. episcopalis sensu stricto (trans α). dmax dph h w u jhc 4061 (ht): mean values: <dmax> 114.3 ± 3.60 σ = 9.53 (8.3%) n = 7 <dph> 70.3 ± 2.28 σ = 8.22 (11.7%) n = 13 <wph> 0.72 ± 0.040 σ = 0.144 (20.0%) n = 13 dmax, dph in mm 0.5172110 0.50 0.11 dmax dph h w u jhc 5158: jhc 5159: ~0.36~6995 0.36 0.18 0.495985 ~0.88 mean values: <dmax> 81.2 ± 1.56 σ = 5.84 (7.18%) n = 14 <dph> 55.67 ± 1.40 σ = 4.85 (8.71%) n = 12 dmax, dph in mm 82 1. the rapidity with which the cranocephalites evolved allows following the microevolution of the ammonites at a level of time-resolution with few rivals in the geo logical record. the 34 transients, bo-1 to po-23, represent evolution over the duration of the late bajocian, which according to the latest geological time scale (gradstein et al. 2012, table 26.3) had an estimated duration of 1.17 myr providing a time resolution here of c. 34 kyr and is close to the achievable limit, set by the ability to recognise morphological changes in successive assemblages. hence, on this time scale, evolution appears to be continuous. 2. the observed record in the rocks continues to be highly discontinuous: the ammonites are found concentrated in thin beds separated by beds that are devoid of them. the biostratigraphic record is therefore almost entirely composed of snap-shots of faunas in time – some beds with ammonites, most beds without. this is in marked contrast to other marine organisms e.g. dinoflagellates, coccoliths, calpionellids, or planktonic foraminifers whose evolutions have been studied biostratigraphically or used for age determinations. 3. discontinuities in the vertical sequences of ammonites are caused by stratigraphic discontinuities rather than by punctuated evolution. geologically, the litho stratigraphic record – the surviving record of sedimentary dynamics, both of input and erosion – is highly incomplete, both vertically and horizontally. as the present account shows, the ammonites reveal the existence of some spectacular non-sequences in the middle jurassic sediments of jameson land whose presence in the rocks is reflected in nothing more than a parting between beds looking like any other. thus, vertically, at trefjord bjerg, the sandstone of horizon po-11 sits directly on the sandstone of bo1: sediments of ages bo-2 up to a level of po-10 are missing. at hurry inlet, sediments of in-1 – j-10 (33 transients) are missing. horizontally, many of the sediments of po-1 – po-3, po-5 – po-9 are lenticular or lost in the region of katedralen in the overlap onto the clinoform unit on a distance scale of kilometres or less. as buckmann (1881) noted, the more complete the biostratigraphical record becomes, the less complete the lithostratigraphic record turns out to be (see callomon 1995, pp. 140, 147). 4. almost all of the ammonite assemblages consist almost exclusively of fully-grown adults and predominantly of macroconchs at that. microconchs are rare. the probable reasons for this point to a closed life cycle in which the fossil accumulations now mark the breeding, spawning and dying grounds of gregarious nektonic organisms that migrated during their life cycles, with sexual segregation in the final stages. acknowledgements the geological institute (university of copenhagen), the geological survey of greenland (ggu, now part of geus, the geological survey of denmark and greenland), the carlsberg foundation, the danish natural research science council, are thanked for support and the opportunity to make the new collections in the ugleelv area. jhc’s field work in jameson land in 1994–1996 was partly funded by norsk hydro and by grants from the danish natural science research council. the visits of jhc to the geological museum in copenhagen were organised by david a.t. harper and funded by eu’s cobice and synthesys museum visitor programmes. carol st. john payne, michael engkilde and michael larsen are thanked for help in the field. paul and peter callomon are thanked for entrusting pal and fs with the material for preparation for publication. carol st. john payne assisted jhc during curation of the collection (in copenhagen) and with photographing fossils for plates. after the death of jhc, carol located notes, electronic files and other material which were invaluable for pal and fs in completing the manuscript. her help is highly appreciated. jette halskov drafted the figures. the referees eckhard mönnig (natur kunde museum, coburg) and paul l. smith (university of british columbia) provided valuable comments on the manuscript. preparation of the manuscript for publication was supported by geocenter denmark. conclusions 83 aldinger, h. 1935: geologische beobachtungen im oberen jura des scoresbysundes (ostgrönland). meddelelser om grønland 99 (1), 128 pp. alsen, p. 2015: description of a new bajocian 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(eds): the jurassic of denmark and greenland and adjacent areas. geological survey of denmark and greenland bulletin 1, 21–59. perch-nielsen, k., henriksen, n. & stemmerik, l. (compilers) 1983: geological map of greenland, 1:100 000, fleming fjord 71 ø.1 nord. copenhagen: geological survey of greenland. piasecki, s., larsen, m., therkelsen, j. & vosgerau, h. 2004a: jurassic dinoflagellate cyst stratigraphy of hold with hope, north-east greenland. in: stemmerik, l. & stouge, s. (eds): the jurassic of north-east greenland. geological survey of denmark and greenland bulletin 5, 73–97. piasecki, s. callomon, j.h. & stemmerik, l. 2004b: jurassic dinoflagellate cyst stratigraphy of store koldewey, north-east greenland. in: stemmerik, l. & stouge, s. (eds): the jurassic of north-east greenland. geological survey of denmark and greenland bulletin 5, 99–112. poulton, t.p. 1978: pre-late oxfordian jurassic biostratigraphy of northern yukon and adjacent northwest territories. in: stelck, c.r. & chatterton, b.d.e. (eds): western and arctic canadian bio-stratigraphy. geological association of canada special paper 18, 445–472. poulton, t.p. & callomon, j.h. 1976: major features of the lower and middle jurassic stratigraphy of northern richardson mountains, northeastern yukon territory and northwestern district of mackenzie. geological survey of canada paper 76-1b, 345–352. poulton, t.p., leskiw, k. & audretsch, a. 1982: stratigraphy and microfossils of the jurassic bug creek group of northern richardson mountains, northern yukon and adjacent northwest territories. geological survey of canada bulletin 325, 137 pp. riccardi a.c. & westermann g.e.g. 1991: middle jurassic am mo noid fauna and biochronology of the argentine-chilean andes. part iii: bathonian–callovian eurycephalitinae, stepha no cera taceae. palaeontographica a216, 1–110. rosenkrantz, a. 1929: preliminary account of the geology of scoresby sound district. in: koch, l.: the geology of east greenland. meddelelser om grønland 73(2), 135–54. rosenkrantz, a. 1934: the lower jurassic rocks of east greenland. part 1. meddelelser om grønland 110(1), 122 pp. salfeld, h. & frebold, h. 1924: juraund kreidefossilien von nowaja semlja. report of the scientific results of the norwegian expedition to novaya zemlya 1921 23, 12 pp. kristiania: viden skapsselskapet i kristiania. schindewolf, o.h. 1925: entwurf einer systematik der peri sphincten. neues jahrbuch für mineralogie, geologie und paläontologie (abteilung b), beilage-band 52, 309–343. sokolov, d.n. 1912: zur ammonitenfauna des petschoraschen jura. mémoires du comité géologique, nouvelle série 76, 1–49 (in russian), 50–65 (german summary). sokolov, d.n. 1913: fossils from the boulders on novaya zemlya. trudy geologicheskogo muzeia imeni petra velikogo impera torskoi akademii nauk 7, 59–92. spath, l.f. 1932: the invertebrate fauna of the bathonian– callovian deposits of jameson land (east greenland). med delelser om grønland 87(7), 158 pp. stauber, h. 1940: stratigraphisch-geologische untersuchungen in der ostgrönländischen senkungszone des nördlichen jameson landes. meddelelser om grønland 114(7), 34 pp. surlyk, f. 1991: sequence stratigraphy of the jurassic – lowermost cretaceous of east greenland. aapg bulletin 75, 1468–1488. surlyk, f. 2003: the jurassic of east greenland: a sedimentary record of thermal subsidence, onset and culmination of rifting. in: ineson, j.r. & surlyk, f. 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(ed.): treatise on invertebrate paleontology, part l, mollusca 4 (revised), volume 4, 362 pp. boulder, colorado: geological society of america & lawrence, kansas: university of kansas press. 86 geological survey of denmark and greenland. bulletin 10, 13-16 the upper cretaceous – danian succession in denmark and most of nw europe is composed mainly of chalk and associated shallower water carbonates deposited in a wide epeiric sea during an overall global sea-level highstand (e.g. surlyk 1997). the maastrichtian–danian chalk has been intensely studied over the last 20 years, since it forms the most important reservoir rock for hydrocarbons in the north sea central graben (e.g. surlyk et al. 2003; klinkby et al. 2005). in denmark, thousands of water wells have been drilled through the succession as about 35% of the water consumption is from maastrichtian chalk and danian bryozoan limestone. during 2005 the new cretaceous research centre (crc) was established jointly at geocenter copenhagen by the geological institute, university of copenhagen and the geological survey of denmark and greenland (geus) with financial support from the danish natural science research council (fnu). crc aims at studying the earth system in a greenhouse world, with special emphasis on the upper cretaceous – danian chalk of nw europe. the stable, longlasting marine macro-environment represented by the chalk sea provides a unique opportunity to analyse and link the depositional, geochemical and biological responses to external forcing at time scales ranging from the sub-milankovitch to the million year range. the studies will be based on a wide range of methods, including seismic stratigraphy, palaeoecology, sequence-, cycloand biostratigraphy, isotope geochemistry, sedimentology and time series analysis. this paper presents the first preliminary results of a crc drilling cam© geus, 2006. geological survey of denmark and greenland bulletin 10, 13–16. available at: www.geus.dk/publications/bull shallow core drilling of the upper cretaceous chalk at stevns klint, denmark lars stemmerik, finn surlyk, kurt klitten, susanne l. rasmussen & niels schovsbo 50 km n < 500 m 500–1000 m 1000–1500 m 1500–2000 m > 2000 m thickness of chalk ringkøbing – danish basin sorgenfrei–tornquist zone outer limit fault salt dome late cretaceous inversion basement high fyn high stevns a r k copenhagen jylland sigerslev højerup st. heddinge rødvig stevns-1 mandehoved stevns-2 sea cliff seismic line village road 1 km b fig. 1. a. thickness of the upper cretaceous – danian chalk in the danish area. r: rørdal quarry; k: karlslunde-1 and tune-1 boreholes. b. map of the study area at stevns showing the location of the stevns-1 and stevns-2 boreholes. 13 paign at stevns klint, eastern denmark (fig. 1), where two shallow boreholes were drilled and logged from near the base of the danian bryozoan limestone and down through the upper 350–450 m of the very thick upper cretaceous chalk section (vejbæk et al. 2003). the cores represent the first complete sections through the maastrichtian chalk of eastern denmark. shallow core drilling seismic data from the offshore area immediately to the east of stevns klint indicate that chalk deposits in this area were modified by powerful, long-lasting bottom currents, and that the late cretaceous sea floor was continuously sculpted by contour-parallel bottom currents into systems of ridges and drifts, moats and valleys with amplitudes up to 150 m and widths of several kilometres (lykke-andersen & surlyk 2004; esmerode et al. in press). in order to investigate the lithological composition of this dynamic chalk system and to provide material for time series analysis, two shallow boreholes were drilled and cored along stevns klint from near the base of the danian bryozoan limestone and 350–450 m down into the upper cretaceous chalk (fig. 1). the northern borehole, stevns-1 (dgu 218.1938) approximately 2 km east of sigerslev was drilled near the culmination of the undulating k–t boundary surface, and is believed to penetrate a ridge succession (fig. 1). the southern borehole, stevns-2 (dgu 218.1945) north-east of rødvig was drilled on a depression in the undulating k–t boundary surface, and was drilled to penetrate a valley-fill succession (figs 1, 2). a comprehensive logging program was subsequently carried out in the two boreholes including spectral gamma, density, sonic, induction, temperature, conductivity, porosity, magnetic susceptibility, resistivity and optic televiewer logs. logging was carried out in open hole using standard methods. stevns-1 was drilled to a depth of 456.1 m with 100% core recovery. it penetrated approximately 12 m of danian bryozoan limestone before entering the maastrichtian chalk succession. preliminary nannofossil data (e. sheldon, personal communication 2005) indicate that the maastrichtian chalk is approximately 400 m thick and that the basal part of the drilled succession is of late campanian age. the drilled succession is roughly divided into an upper campanian – lowermost maastrichtian interval of bioturbated chalk with rare thin clay beds that gradually passes up into a 50 m thick succession of interbedded lower maastrichtian chalk and marl. the gamma ray log indicates that marl layers become gradu14 fig. 2. the fakse kalk a/s diamant board 747 wireline rig at the site of stevns-2 in the abandoned boesdal quarry north-east of rødvig. metres total gamma (cps)5 30 -100 -200 -300 -400 fig. 3. gamma-ray log of stevns-1 showing the presence of two marl-rich intervals from c. 65–100 m and 345–370 m. ally less abundant upwards, and the interval from 300–100 m is composed of almost pure chalk (fig. 3). the first flint nodules and layers appear at 135 m and flint is common in the upper part of the succession. the 100–70 m interval consists of flint-rich chalk with distinctive marly horizons while the uppermost part of the maastrichtian succession shows an increasing content of bryozoans and other macrofossils. stevns-2 was drilled to a depth of 350 m with 100% core recovery. it penetrated 4 m of danian limestone and approximately 300 m of maastrichtian chalk before it terminated in upper campanian chalk. the drilled succession shows the same overall changes in lithology as stevns-1 but the thickness of the individual units seem to vary considerably. this is particularly true for the chalk-dominated succession between the two marly intervals which thins from approximately 200 m in stevns-1 to 120–130 m in stevns-2. core scanning the 456.1 m long stevns-1 core has been scanned at the geus core laboratory using a set-up which allows simultaneous spectral gamma-ray and density measurements (fig. 4). the spectral gamma-ray analysis is carried out using two 15 cm nai (tl) crystals and the bulk density is determined using a caesium source. the scanning was performed using a speed of 1 cm/min., corresponding to a vertical resolution of approximately 2 cm for the density log. the scanning data thus supply high resolution data to support sedimentary and geochemical data from the core, as exemplified in fig. 4 illustrating the cyclic nature of the succession. in this interval uranium only shows minor variations and the high total gamma readings correspond to kand th-peaks. it is thus evident that these peaks correspond to marly beds with high levels of clay-bound k and th whereas intervals with low total gamma readings correspond to purer chalk. comparison with other danish chalk sections the stevns-1 and stevns-2 cores represent the first complete sections through the maastrichtian chalk of eastern denmark. the uppermost up to 35 m of the maastrichtian are exposed along stevns klint and in quarries along the cliff. further south, on the island of møn, approximately 70 m of lower maastrichtian chalk occur as glacially disturbed thrust sheets in excellent cliff exposures dominated by flint-rich bioturbated chalk with rare incipient hardgrounds; at stevns klint the youngest maastrichtian deposits become gradually richer in bryozoans. the main difference between the cored section and the adjacent outcrops is the lack of flint in the deeper part of the core. the danian – uppermost maastrichtian succession has been drilled by numerous water wells in the greater copenhagen area to the north and north-west of stevns klint, and more recently the upper 250 m of the maastrichtian chalk were cored in tune-1 (dgu 207.3841) and karlslunde-1 (dgu 207.3850; fig. 1; larsen et al. 2006). these two cores show a facies development similar to that seen in the upper part of the stevns cores with a downward change from bryozoan-rich chalk to bioturbated chert-rich chalk interrupted in the mid-maastrichtian by a succession of interbedded marl 15 , metres total gamma (cps)0 2 k (%)-0.1 0.6 u (ppm)0 5 th (ppm)0 6 -345 -347.5 -350 -352.5 -355 fig. 4. core scan log of the 345–355 m interval in stevns-1 showing total gamma response and k, u and th concentrations. note that intervals with high total gamma correlate with high concentrations of k and th, and correspond to marly beds. and chalk layers before the core bottoms in midmaastrichtian chalk (larsen et al. 2006). the most important difference is the presence of a distinctive 20 cm thick marl, interpreted as representing the kjølby gaard marl of troelsen (1937, 1955), 9–10 m below the k–t boundary in tune-1. this marl has also been reported from adjacent water wells (larsen 1997) and in quarries in northern jylland (troelsen 1955), and seems to be a distinctive stratigraphic marker bed in northern denmark. the mid-maastrichtian interval of interbedded chalk and marl drilled in stevns-1, stevns-2 and karlstrup-1 also seems to represent a distinctive regional event as a similar cyclic chalk-marl succession is known from the rørdal quarry in northern jylland. future work the stevns cores represent the first continuous sections of the maastrichtian chalk in nw europe and provide a unique possibility for sedimentological, geochemical and isotope geochemical investigations as well as time series analysis and studies of reservoir properties. the stevns area has been buried to less than 600–700 m during post-danian time and the lack of burial diagenetic overprinting makes the carbonates ideal for isotope geochemical analyses to monitor both the global carbon cycle and to provide information on the temperature and salinity of the late cretaceous boreal ocean. core data will be integrated with log data to better understand the lithological significance of the log responses, and 2d depositional models of the chalk will be based on integration of core and log data with reflection seismic data along a line connecting the two boreholes. reflection seismic data have been collected between the two boreholes in cooperation with holger lykke-andersen, university of aarhus, and refraction seismic data have been collected by lars nielsen, university of copenhagen. the two data sets are still in the stage of processing but will provide important 2d data to support the information from the cores. acknowledgements the danish natural science research council is thanked for financial support. additional funding for logging has been obtained from geocenter copenhagen. fakse kalk a/s and rambøll a/s provided technical assistance. stevns kommune and stevns natur center kindly allowed us to drill at their property. references esmerode, e.v., lykke-andersen, h. & surlyk, f. in press: ridge and valley systems in the upper cretaceous chalk of the danish basin: contourites in an epeiric sea. geological society of london, special issue. klinkby, l., kristensen, l., nielsen, e.b., zinck-jørgensen, k. & stemmerik, l. 2005: geological characterisation of the kraka field chalk reservoir, danish north sea – integration of seismic and log data. petroleum geoscience 11, 113–124. larsen, f., sonnenborg, t.o., madsen, p., ulbak, k.a. & klitten, k. 2006: saltvandsgrænsen i kalkmagasinerne i nordøstsjælland; delrapport 6: saltvandsudvaskning i danienkalk og skrivekridt – detailundersøgelser i karlslunde værkstedsområde. danmarks og grønlands geologiske undersøgelse rapport 2006/21, 103 pp. larsen, o. 1997: mapping of the maastrichtian–danian boundary in the coastal area of køge bugt by gamma and resistivity logging. bulletin of the geological society of denmark 44, 101–113. lykke-andersen, h. & surlyk, f. 2004: topography of the cretaceouspalaeogene boundary at stevns klint, denmark: inversion tectonics or primary relief of the chalk sea-floor? journal of the geological society (london) 161, 343–352. surlyk, f. 1997: a cool-water carbonate ramp with bryozoan mounds: late cretaceous – danian of the danish basin. in: james, n.p. & clarke, j.a.d. (eds): cool-water carbonates. sepm (society for sedimentary geology) special publication 56, 293–307. surlyk, f., dons, t., clausen, c.k. & higham, j. 2003: upper cretaceous. in: evans, d., graham, c., armour, a. & bathurst, p. (eds/coordinators): the millennium atlas: petroleum geology of the central and northern north sea, 213–233. london: geological society. troelsen, j.c. 1937: om den stratigrafiske inddeling af skrivekridtet i danmark. meddelelser fra dansk geologisk forening. bulletin of the geological society of denmark 9, 260–263. troelsen, j.c. 1955: globotruncana in the white chalk of denmark. micropaleontology 1, 76–82. vejbæk, o.v., bidstrup, t., britze, p., erlstrøm, e., rasmussen, e.s. & sivhed, u. 2003: chalk structure map of the central and eastern north sea. danmarks og grønlands geologiske undersøgelse rapport 2003/106, 55 pp. 16 authors’ addresses l.s., s.l.r., k.k. & n.s., geological survey of denmark and greenland, geocenter copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ls@geus.dk f.s., geological institute, university of copenhagen, geocenter copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. geological survey of denmark and greenland bulletin 4, 2003, pp 29-32 29 groundwater is the major source of drinking water in many european countries, and in denmark alone it accounts for more than 99% of the drinking water supply. within the past decade pesticide residues have frequently been detected in groundwater, in many cases at levels exceeding the 0.1 µg/l limit set by the european community. as a consequence, drinking water abstraction wells have had to be closed in many places in denmark and other european countries, and a vast amount of money is expended to monitor groundwater pesticide levels. a degradation product of the herbicide dichlobenil, 2,6-dichlorobenzamide (bam), is the most common cause of drinking water well closure in denmark. triazines and their metabolites also contaminate groundwater in many countries, and pose a similar risk to the drinking water supply. analysis of most pesticides and their degradation products is usually carried out by concentrating the samples by solvent extraction, and identifying the contaminants using gas chromatography (gc) or high-pressure liquid chromatography (hplc) combined with mass spectrometry (ms). these methods, although robust and well established, are very time-consuming and require specialised instrumentation. the large quantity of solvents used is another drawback to these methods, as the solvents themselves may be carcinogenic and are also well known contaminants of groundwater. the development of cheap, more sensitive and more rapid pesticide assays is therefore urgent. due to their very high sensitivity, immunological methods have long been used in biological science for analysing a large variety of organic structures, but have only recently been introduced to environmental analysis. the benefit of such assays is primarily their high sensitivity, which allows the analysis to be undertaken without the need to concentrate the samples, but also the facility of dealing with large numbers of samples. compared to conventional analyses, immunological methods face two major drawbacks – one related to specificity and the other to the fact that only very few chemicals can currently be analysed simultaneously. the crux of the specificity problem is that although antibodies react very specifically with particular chemical structures, these same structures may be present in analogous compounds. thus antibodies developed to recognise, for example the herbicide atrazine might also recognise other triazines (bruun et al. 2001). an important scientific challenge is therefore the development of highly specific assays recognisimmunological analysis of pesticides: a new tool in groundwater testing jens aamand, leif bruun and claus bo vöge christensen fig. 1. development of monoclonal pesticide antibodies is initiated by covalent conjugation of the pesticide to a carrier protein. this pesticide-carrier complex is injected into mice and after approximately two months the mice have produced antibodies against the pesticide. selected mice are sacrificed and the spleen is removed to isolate the antibody-producing cells. these cells are difficult to cultivate in vitro, and they are therefore fused with myeloma cells to ensure the viability of the antibody-producing cells. the fused hybridoma cells are cultivated, tested and isolated to achieve monoclonal cultures, which produce one type of antibody only with special characteristics such as binding efficiency and specificity. slightly modified from aamand et al. (2003). geological survey of denmark and greenland bulletin 4, 29–32 (2004) © geus, 2004 ing each individual compound, as well as assays recognising groups of related chemicals. with respect to the simultaneous analysis of numerous chemicals, this can be resolved by implementing the new biochip technology, which incorporates the parallellity of sample screening. on a pesticide biochip many specific immunological assays are carried out in isolated small spots on a glass or polymer surface. each spot has a size of approximately 150 micrometers and forms a specific analysis. such a miniaturised platform will be usable for monitoring programmes where water samples have to be screened for a range of chemical contaminants. the overall objectives of this study have been (1) to develop immunoassays for high-sensitivity analysis of specific pesticides and chemically related groups of pesticides, and (2) to transfer the developed assays to a miniaturised biochip platform in a manner allowing analysis of several pesticides simultaneously. immunological analysis of pesticides the basis for the development of new immunological analyses is the antibody that reacts with complementary molecules, the so-called antigens. antibodies are part of the immunological defence system in animals and humans for protection against pathogenic vira and bacteria. following an infection, the organism produces antibodies that recognise and bind to specific molecular structures on the surface of the penetrating bacteria or virus (the antigens). upon the binding of antibodies, other effector functions of the immune system identify and destroy the bacteria. the chemical structures of the pesticides themselves are too small to induce an immunological response. however, by linking the pesticides to larger carrier molecules it is possible to deceive the immune system into starting the production of antibodies against the pesticide (fig. 1). to initiate antibody production the pesticide-carrier complex is injected into an animal, e.g. a rabbit or a mouse, thus inducing an immunological response resulting in the production of antibodies against the pesticide-carrier complex. antibodies are produced by so-called b-cells each producing a single antibody species, which recognise a specific structure on the pesticide. as the animal contains many b-cells which all produce antibodies, a range of antibodies reacting with different structures on the pesticide and with different affinity will be generated. by purification of the antibodies from the blood serum, a polyclonal antibody serum is obtained, containing antibodies from different b-cells. however, it is often more appropriate to produce monoclonal antibodies (mab), i.e. specific antibodies all arising from the same b-cell clone. production of monoclonal antibodies when it has been established by serological screening that the immunised animal produces antibodies with the correct specificity, the spleen, which contains many antibody-producing cells, is removed and grown in culture. clones producing antibodies with the desired properties are then selected. it is possible to select antibodies that react with chemical structures specific for a single pesticide molecule, or alternatively a structure shared by a group of pesticides such as the triazines. in addition to their high specificity monoclonal antibodies also have the advantage of consistency. it is always possible to reculture the hybridomas and produce further antibodies with exactly the same characteristics. development of immunological assays the next step following the selection of suited antibodies is the development of an immunochemical pesticide assay. the analysis is often carried out as a so-called competitive immunoassay in microtitre plates, which are preformed plastic plates with 96 wells (fig. 1). a known amount of pesticide is immobilised on the bottom of the wells. the samples to be analysed are added to the wells (typically 100 µl), followed by 30 fig. 2. example of a competitive assay. in cases with high concentrations of pesticides in the sample, fewer antibodies will bind to the pesticide immobilised to the surface (e.g. the bottom of the microtitre plate) and a low signal will be measured. if the samples do not contain the pesticide, maximal amounts of antibody will bind to the surface resulting in a high signal. slightly modified from aamand et al. (2003). the addition of the antibodies. at this point the antibodies can react with either the pesticide immobilised in the microtitre plates or the pesticide in the sample itself. if the concentration of the pesticide in the sample is low, more antibodies will react with the immobilised pesticides and vice versa. when the reaction is completed the microtitre plate wells are washed, leaving only the immobilised pesticideantibody complex (fig. 2). the antibodies can be directly monitored if coupled with an enzyme, catalysing an enzymesubstrate reaction that yields a coloured end product. the accumulation of the end product is then monitored by absorbance measurements and compared spectrophotometrically to the absorbance of known standards (fig. 3). new immunological assays have been developed for several triazines including their degradation products (bruun et al. 2000a, b, 2001) and for bam (bruun et al. 2000c). all assays have a very low detection limit in the range of 0.01–0.02 µg/l, making them ideal for monitoring specific pesticide residues in groundand drinking water. analyses are typically carried out in four replicates, and each microtitre plate also contains a number of pesticide standards. a total of 13 samples can be analysed at each microtitre plate within a period of 3–4 hours. from microtitre plate to pesticide biochip one of the drawbacks of the microtitre format is that it is only possible to analyse for one pesticide in each routine. however, changing the analysis format from microtitre plates to biochips allows for the analysis of several compounds simultaneously. the term ‘biochips’ describes an analysis where the chemical reactions are not separated by wells, but are carried out on a planar surface such as a glass slide. the principle of the analysis is the same as for the microtitre plates, but the reagents are added as microspots (in the nanoliter range) on the glass surface. using a robot equipped with printing pins, about 2000 samples/cm2 can be added as separate spots on the surface. all chemical reactions are then carried out within the individual spots. for the pesticide biochip fluorescence-conjugated monoclonal antibodies were used, which enables the detection of separate signals from each spot on the surface by use of a laser scanner (fig. 4). as a result of the small dimensions, the individual reactions equilibrate faster and the complete analysis of a biochip can be carried out within 90 minutes. the analyses on the biochips are also more sensitive than on microtitre plates. we have developed a pesticide biochip for bam and atrazine with a sensitivity of about 1 ng/l (fig. 5), which is 100 times less than the limit value for drinking water set by the eu (belleville et al. 2003, 2004). possibilities and limitations the benefits of the immunochemical analyses compared to chromatographic techniques are that: (1) less sample volume is needed, which means an easier transport of samples to the laboratory; (2) no solvents or other chemicals are necessary which potentially could pollute the environment; and (3) the immunochemical analyses are much cheaper to carry out. at present, the immunochemical techniques only enable the analysis of a few compounds simultaneously. in contrast, chromatographic methods (e.g. hplc or gc/ms) provide the concentration of a range of compounds within the same routine. however, use of the pesticide biochips opens the possibility of analysis of more compounds simultaneously. at present the biochip includes bam and atrazine only, but in theory it is possible to include additional pesticides as soon as usable antibodies become available. another problem to be faced is related to the specificity of the antibodies. many antibodies may react not only with the targeted pesticide, but also with chemically related compounds. this is the case for the atrazine antibodies that may also react with other triazine herbicides (bruun et al. 2001). the specificity, however, is not a problem with bam, because the reactivity of the antibody with other compounds is negligible (bruun et al. 2000c). the new pesticide biochip enables the analysis of pesticides in a single drop of water in concentrations as low as 1 ng/l. in principle the pesticide biochip allows the analysis of a range of pesticides, but for the development of such multicomponent analysis further antibodies are needed with high specificities to the individual pesticides. 31 fig. 3. example of a standard curve. note the inverse relationship between pesticide concentration and signal. slightly modified from aamand et al. (2003). 32 acknowledgement the present work is supported by the immunalyse project (grant no. 9901188) financed by the danish research agency. references aamand, j., bruun, l. & christensen, c.b.v. 2003: mus hjælper til med pesticidanalyser. dansk kemi 84, 29–31. belleville, e., dufva, m., aamand, j., bruun, l. & christensen, c.b.v. 2003: quantitative assessment of factors affecting the sensitivity of a competitive immunomicroarray for pesticide detection. biotechniques 35, 1044–1051. belleville, e., dufva, m., aamand, j., bruun, l., clausen, l. & christensen, c.b.v. 2004: quantitative microarray pesticide analysis. journal of immunological methods 286, 219–229. bruun, l., koch, c., jakobsen, m.h. & aamand, j. 2000a: a new monoclonal antibody for the sensitive detection of cyanazine and other striazines in water by elisa. food and agricultural immunology 12(4), 253–262. bruun, l., koch, c., jakobsen, m.h. & aamand, j. 2000b: new monoclonal antibody for the sensitive detection of hydroxy-s-triazines in water by enzyme-linked immunosorbent assay. analytica chimica acta 423, 205–213. bruun, l., koch, c., pedersen, b., jakobsen, m.h. & aamand, j. 2000c: a quantitative enzyme-linked immunoassay for the detection of 2,6dichlorobenzamide (bam): a degradation product of the herbicide dichlobenil. journal of immunological methods 240, 133–142. bruun, l., koch, c., jakobsen, m.h., pedersen, b., christiansen, m. & aamand, j. 2001: characterisation of monoclonal antibodies raised against different structures belonging to the s-triazine-group of herbicides. analytica chimica acta 436, 87–101. fig. 4. laserscan of a pesticide biochip designed to analyse for bam and atrazine. each spot represents a single analysis of a standard with a known concentration. within each concentration the six spots to the left are bam and the six spots to the right are atrazine. slightly modified from aamand et al. (2003). authors’ addresses j.aa., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: jeaa@geus.dk l.b., statens serum institut, artillerivej 5, dk2300 copenhagen s, denmark. c.b.v.c., technical university of denmark, department of microand nanotechnology, ørsted plads 345, dk-2800 kongens lyngby, denmark. << /ascii85encodepages false /allowtransparency false /autopositionepsfiles true /autorotatepages /all /binding /left /calgrayprofile (dot gain 20%) /calrgbprofile (srgb iec61966-2.1) /calcmykprofile (u.s. web coated \050swop\051 v2) /srgbprofile (srgb iec61966-2.1) /cannotembedfontpolicy /warning /compatibilitylevel 1.4 /compressobjects /tags /compresspages true /convertimagestoindexed true /passthroughjpegimages true /createjdffile false /createjobticket false /defaultrenderingintent /default /detectblends true /colorconversionstrategy /leavecolorunchanged /dothumbnails false /embedallfonts true /embedjoboptions true 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<feff4f7f752890194e9b8a2d5b9a5efa7acb76840020005000440046002065874ef65305542b8f039ad876845f7150cf89e367905ea6ff0c4fbf65bc63d066075217537054c18cea3002005000440046002065874ef653ef4ee54f7f75280020004100630072006f0062006100740020548c002000520065006100640065007200200035002e0030002053ca66f465b07248672c4f86958b555f3002> >> >> setdistillerparams << /hwresolution [2400 2400] /pagesize [595.000 842.000] >> setpagedevice geological survey of denmark and greenland. bulletin 10, 21-24 the national geological database at the geological survey of denmark and greenland (geus) is based on an extensive well database jupiter, a geophysical database gerda (tulstrup 2003) and a recently established database for various types of geological models. these databases are integrated in a gis system. the integration of this data enables new possibilities of constructing improved geological models. gis systems offer a powerful tool for the geologist not only in combining multiple data, but also in visualising the model and hence presenting the final product in a simple and understandable way. 3d geological models will become increasingly important for the execution of improved cost-benefit analysis and risk assessment of contaminated sites, as well as strategic evaluation of groundwater and raw material resources in general. the possibility of storing such models on a public platform will be a major advance for future users of geological databases. the primary goal of this paper is to demonstrate the potential of an integrated gis system, with an example of how traditional geological information may be combined in new ways in order to improve the correlation of well data in multiple directions. the application is demonstrated for a highly contaminated industrial site in the town of ringe, denmark (fig. 1). data analysis basically two methods exist for constructing 3d models from borehole data. using the vertical approach, information from the nearest boreholes is projected onto an array of vertical sections (fig. 2), thus guiding the drawing of lines separating the interpreted units. the boundaries separating the geological units from all cross-sections are then converted into a 3d surface through interpolation. using the horizontal approach, information from all boreholes is projected onto an array of horizontal sections covering the relevant area (fig. 3). the construction of 3d geological models in glacial deposits to characterise migration of pollution knud e.s. klint, frants von platen-hallermund and mette christophersen © geus, 2006. geological survey of denmark and greenland bulletin 10, 21–24. available at: www.geus.dk/publications/bull 21 3d model 600 m ringe denmark n rta site possible plume area source area a b wells cross section 50 m 1322 1321 1320 1319 1318 1317 1305 1304 1303 1302 1301 1300 1299 1298 1065 1064 1063 1062 1061 1056 1055 1054 1053 1051 1017 1016 1015 1014 1013 1011 1009 998 968 957 949 946 941 936 935 933 930 924 922 921 872 859 855 854 852 851 850 822 821 820 819 816 812 811 810 795 794 793 792 122 121 108 27j fig. 1. location of the former tar and asphalt factory in ringe (rta site) with the distribution of wells inside the source area and in a wider zone around the site. the site is situated on map sheet 155, and all well numbers have the prefix ‘155.’ in the jupiter well database. vertical section a–b shown in fig. 2; a horizontal slice diagram and 3d model are shown in figs 3–4. distance between the sections is determined by the geological complexity. the lithological and hydraulic information from the boreholes penetrating the sections is used to guide drawing of unit boundaries at the relevant level. finally, the 3d model is constructed by stacking all the horizontal sections (fig. 4). in this study both methods are used, but with emphasis on the horizontal sections because current knowledge of the glaciodynamic history and position of the local geological units are better utilised in this way. a general geomorphological analysis and thorough description of the depositional environment is included in the interpretation of the different geological units in the area (klint & von platen-hallermund 2006). database procedures for extracting borehole data from a series of horizons, as well as gis procedures for producing a collection of maps based on borehole information or interpreted unit boundaries in the same intervals have been developed as a tool for the construction of improved geological models. the 3d display of the gis system is used for viewing the final model. site history the ringe tar and asphalt factory site (rta site) is located in the centre of ringe (fig. 1). asphalt and creosote production took place from 1929 to 1962, and during that period storage tanks were placed all over the area. in 1987 it was discovered that the factory site subsurface was strongly contaminated with creosote, and in 1988 remediation of the site was commenced. old buildings and storage tanks were removed and some of the most heavily contaminated soil was excavated and removed. the site has been used extensively for contamination investigations for several years (e.g. broholm et al. 2000; klint & tsakiroglou 2000). a large number of wells have been drilled during these projects, and four open pits were excavated. 22 108 1318(f4) 1317(f3) 1063 794 1056 1015 1319(f5) 50 100 150 200 1064 1298(k1) 1322(k9) 816 792 1321(k8) clay till sandy till glaciofluvial silt/clay glaciofluvial sand glaciofluvial gravel fill ground water level contamination unit 8 unit 7 unit 7 unit 4 unit 3 ? ? unit 2 unit 1 unit 6 unit 5 250m unit 1 unit 4 unit 3 unit 2 a b m a b o ve s ea l ev el unit 3 unit 6 75 70 65 60 55 50 45 40 35 30 80 kote 61 m above sea level 60 m above sea level unit 4 basal clay till unit 4 basal clay till unit 3 braided river system clay sand gravel wellsareas vertical crosssection a-b 100 m fig. 2. vertical correlation approach. crosssection a–b (fig. 1) showing the distribution of the eight primary geological units. three new wells display grain size distributions in grey shading. all well numbers are prefixed ‘155.’. fig. 3. horizontal correlation approach. slice diagram showing the distribution of clay, sand and gravel in wells at two depth intervals. red frame indicates location of the ringe tar and asphalt factory site. position of section a–b in fig. 2 is indicated. geological framework the ringe tar and asphalt factory site is situated on top of an elongated hill approximately 80 m above sea level. several small hills and depressions without runoff indicate that the sediments in the hilly area were probably deposited from a stagnating glacier (dead-ice relief ) and overridden during a later ice advance. during the earlier research projects a local geological model was established for the upper 22 m of the ringe site (sidle et al. 1998, nilsson et al. 2001). the investigations showed that the site is dominated by sediments deposited during the late weichselian glaciation 25 000 – 11 500 years ago. during 2005 nine wells were drilled in the source area to c. 20 m below ground surface, and five deep wells were drilled downstream for locating the contamination plume (fig. 2). the wells were described in terms of matrix texture, structures, colour, fractures, consolidation and contamination. hydraulic tests and groundwater table measurements in both secondary and primary aquifers supported the interpretation of the depositional environment. the conceptual geological model is well constrained near the surface, but becomes more speculative at greater depth due to the decreasing amount of well data. the eight geological units described below were deduced from the cross-sections and log descriptions (fig. 2), and slice diagrams produced with a 1 m interval from 26 to 80 m above sea level (fig 3). finally, a 3d geological model was constructed to outline the spatial distribution of the geological units (fig. 4). unit 1 is basal till. the upper 4–8 m are dominated by a continuous, massive, undulating ground moraine with occasional minor sand lenses which covers most of the area. this till may be classified as a basal till deposited below a glacier that transgressed the area from the east-south-east during the c. 16 kyr young baltic ice advance (nilsson et al. 2001; houmark-nielsen & kjær 2003). the till is penetrated by fractures, and there is generally good hydraulic contact to the underlying beds through the fracture network and embedded sand lenses. unit 2 comprises mixed fluvial and diamict deposits. from 4–16 m below the ground surface, a heterogeneous glacial complex dominated by glaciofluvial silty, sandy and gravelly deposits is interbedded with clayey and sandy diamict sediments. the sediments are partly deformed by glaciotectonic processes, and isolated sand lenses form small secondary water reservoirs especially in the central part of the rta site, at c. 8–9 m below ground surface. other lenses are unsaturated and hence in hydraulic contact with the primary aquifer. approximately 12 m below ground surface a more widespread, 3–4 m thick layer of unconsolidated clayey till covers most of the area, though it may be mixed with larger sand bodies in places. the till is fractured locally and contains numerous thin sand lenses. it is classified predominantly as flow till. the general topography in the area includes several small depressions with the characteristics of dead-ice holes, and the whole unit is interpreted to represent a dead-ice landscape overridden by a glacier. unit 3 consists of braided river valley deposits. between 14–21 m depth a widespread sand/gravel layer 2–7 m thick covers most of the area and probably represents a braided river system with minor channels eroded into an underlying ground moraine. the channels are generally dominated by coarse glaciofluvial sand/gravel and boulders. unit 3 truncates the underlying clay till (unit 4) locally, thus creating a good hydraulic contact with a major sandy aquifer below, which dominates the western and northern parts of the site (unit 5). unit 3 is partly saturated, as the water table is here located approximately 18 m below ground surface. accordingly, the deepest, saturated parts of the channels act as hydraulic avenues for infiltration of groundwater and contaminants from the upper units. the general groundwater and contamination flow is directed towards the west-south-west at the rta site, but turns southwards approximately 30 m west of the site, thus indicating preferential flow controlled by channels (fig. 1). 23 n depth interval 26–61 m above sea level depth interval 26–40 m above sea level unit 4 unit 3 unit 6 unit 6 unit 8 unit 8 unit 8 500 m depth interval 26–80 m above sea level unit 8 50 m 40 unit 7 unit 7 unit 7 300 m unit 6 30 unit 5 unit 5 unit 4 20 unit 4 unit 3 unit 1 unit 2 10 fig. 4. 3d geological model of the ringe site showing the distribution of clayey deposits (left) and sandy deposits (right) at different depth intervals (note vertical exaggeration). ground surface with buildings and position of wells shown in lower model. 24 unit 4 is basal till. a massive basal till 4–8 m thick covers most of the south-eastern area 17–22 m below ground surface. its surface is situated above groundwater level in some areas. it is completely eroded by unit 3 just west of the rta site, where unit 3 is resting directly on unit 5 (fig. 3). unit 5 consists of sandy and silty melt-water river deposits. unit 5 constitutes a more widely distributed occurrence of glaciofluvial sand, occurring between 24–45 m depth, with a generally finer grain size than unit 3. the sand becomes very fine in the lowest parts, and is dominated by glaciofluvial silt and clay below 37–40 m depth. unit 5 is deeply incised into the underlying units, and is thought to represent a well-defined river valley filled with generally finer material than unit 3. the contaminant plume seems to follow this unit in a narrow south-directed fan. unit 6 is clay till. a thin clay till unit locally separates the two sandy units 5 and 7 (fig 4). this unit probably represents erosional remnants of a previously more widespread clay till unit located 33–36 m below ground surface. the lithological and structural information is, however, sparse and any interpretation of its origin is therefore highly speculative. unit 7 comprises river valley deposits. at least 14 m of sand and gravel have been encountered in wells 155.108 and 155.1318 between 35–49 m below ground surface (fig. 2). the water-supply well 155.108 has produced 45 m3 of water per hour and is considered to be located within an extensive sand/gravel body representing a major buried river valley incised in the clay till of unit 8. well 155.1318 has much smaller capacity and hence must be related to a somehow smaller sand reservoir although hydraulically connected to 155.108. unit 8 is basal till. generally little information exists on the lowest clay till unit below 40 m depth. the till is generally massive and well consolidated, with a medium to high content of boulders. it has the nature and appearance of a basal till and is locally more than 10 m thick. only six wells reach this unit, and its distribution is consequently largely unknown. conclusions the combination of correlating multiple geological horizontal sections with traditional vertical geological cross-sections using gis has facilitated the construction of a detailed 3d geological model of the contaminated tar and asphalt factory site in ringe, which outlines the spatial distribution of primarily sandy and clayey sediments. detailed descriptions of the sediment properties were used to interpret the depositional processes and hence the depositional environment. this interpretation was included in the separation of the different deposits into units related to distinct glacial environments. the model will be used for assessing the risk of polluting nearby groundwater reservoirs and for designing an optimal remediation programme. acknowledgements the work was financed by fyns amt and carried out in cooperation with orbicon. references broholm, m.m., rügge, k., tuxen, n., mosbæk, h. & bjerg, p.l. 2000: migration and degradation of pesticides in an aerobic groundwater aquifer: field injection experiments. in: bjerg, p.l., engesgaard, p. & krom, t.d. (eds): proceedings of the international conference on groundwater research, copenhagen, denmark, 6–8 june, 2000, 169–170. rotterdam: balkema. houmark-nielsen, m. & kjær, k.h. 2003: southwest scandinavia, 40–15 kyr bp: palaeogeography and environmental changes. journal of quaternary science 18, 765–786. klint, k.e.s. & tsakiroglou, c.d. 2000: a new method of fracture aperture characterisation. in: tsihrintzis, v.a. et al. (eds): proceedings of the 5th international conference on restoration and protection of the environment 1, thassos, greece, 3–6 july, 2000, 127–136. klint, k.e.s. & von platen-hallermund, f. 2006: geologisk model af rtagrunden ved villavej i ringe. en 3-d geologisk model af kildeområde og faneområde ved den tidligere ringe tjære asfalt fabrik på villavej i ringe. danmarks og grønlands geologiske undersøgelse rapport 2006/9, 18 pp. nilsson, b., sidle, r.c., klint, k.e.s., bøggild, c.e. & broholm, k. 2001: mass transport and scale-dependent hydraulic tests in a heterogeneous glacial till – sandy aquifer system. journal of hydrology 243, 162–179. sidle, r.c., nilsson, b., hansen, m. & fredericia, j. 1998: spatially varying hydraulic and solute transport characteristics of a fractured till determined by field tracer tests, funen, denmark. water resources research 34, 2515–2527. tulstrup, j. 2003: environmental data and the internet: openness and digital data management. geological survey of denmark and greenland bulletin 4, 45–48. authors’ addresses k.e.s.k. & f.v.p.-h., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark, e-mail: kesk@geus.dk m.c., fyns amt, ørbækvej 100, dk-5220 odense sø, denmark. geological survey of denmark and greenland bulletin 4, 2003, pp 17-20 17 the geological survey of denmark and greenland (geus) has for many years been involved with research, advisory and consultancy services concerning the assessment of the geothermal energy potential in denmark, in close cooperation with private and public partners. the survey’s particular responsibility has been the development of geological models to describe and predict the distribution of sandstone reservoirs suitable for geothermal exploitation. danish geothermal resources in known sandstone aquifers are estimated to be sufficient to cover household heating requirements in denmark for more than a century (sørensen et al. 1998). background utilisation of geothermal energy is a well-established technology with more than one hundred plants currently operating in europe. concerns with respect to co2 emission to the global atmosphere have led to increased interest in the utilisation of geothermal energy as one possible way of reducing the consumption of fossil fuels. in 1983, dansk olie & naturgas a/s (dong a/s) was granted a sole concession for the exploration and production of geothermal energy in the entire land area in denmark. in fig. 1. map of denmark showing the regional geothermal potential of possible aquifer formations, based on a burial depth of 1000–2500 m and a sand thickness of more than 25 m. white areas in denmark indicate that the reservoir is not present (ringkøbing–fyn high), too shallow (northernmost jutland), or too deeply buried (central part of danish basin). the locations of the thisted geothermal plant and the new geothermal site at margretheholm in copenhagen are shown. geological survey of denmark and greenland bulletin 4, 17–20 (2004) © geus, 2004 geothermal energy in denmark lars henrik nielsen, anders mathiesen and torben bidstrup 1993 and 2003, selected parts of the concession area were returned to the state in accordance with the licensing terms. the first comprehensive study of danish geothermal resources was presented by michelsen et al. (1981), which incorporates seismic and well data from the danish onshore areas, with a focus on sandstone aquifers between 2000 and 3000 m depth. the survey’s contributions to the two volumes of the atlas of geothermal resources published by the euro-pean commission (haenel & staroste 1988; hurter & haenel 2002) have presented data and information that identify areas of interest for further geothermal exploration. initially it was believed that geothermal heat could be produced from deep, hot aquifers, and in the early 1980s sandstones of the upper triassic gassum formation were tested at depths of c. 3000 m in three wells in northern jutland (fig. 1). thick sandstones were encountered, but permeability was insufficient and the results were discouraging. however, technological innovation during the past decade has shifted interest from deep and hot, but high-risk reservoirs, towards shallower aquifers with good porosity and permeability and thus the potential of producing large volumes of warm water. following heat extraction, the cold water is re-injected into the aquifer at some distance from the producing well via an injection well, in order to maintain reservoir pressure and avoid mixing the cold return water with the warm formation water. areas of potential interest four major structural features – the danish basin, the sorgenfrei–tornquist zone, the ringkøbing–fyn high and the north german basin – exert the overall control on the geothermal prospectivity of denmark. they essentially determine the distribution, thickness, facies types and burial depths of the stratigraphic units with potential reservoirs (fig. 1). the danish basin is bounded by the ringkøbing–fyn high to the south and the sorgenfrei–tornquist zone to the north-east. the upper permian – cenozoic basin-fill is 5–6.5 km thick along the basin axis, increasing locally to more than 9 km in the sorgenfrei–tornquist zone. the triassic – lower cretaceous succession has a relatively uniform thickness in most of the basin with some thinning towards the ringkøbing–fyn high. due to uplift of most of the basin and the ringkøbing–fyn high in early middle jurassic time, the triassic – lower jurassic succession is truncated by the ‘base middle jurassic unconformity’, which shows a progressively deeper truncation towards the ringkøbing–fyn high (fig. 2). on the high the lower jurassic, and in places parts of the triassic, have been eroded. regional subsidence gradually took over again in late middle – early late jurassic time and became more widespread, as shown by a progressively younger upper jurassic – lower cretaceous onlap onto the unconformity towards the high. these events have great influence on the distribution of reservoirs and the thickness of the overburden. thus, the lower triassic reservoirs may be found at moderate depths on the ringkøbing–fyn high and along the northern and southern (north german basin) flanks of the high. the sorgenfrei–tornquist zone crosses northern jutland, kattegat, the northern part of øresund and southern sweden. it is a strongly block-faulted zone with tilted palaeozoic fault blocks overlain by thick mesozoic deposits (fig. 1). this zone experienced continuous, but slow, subsidence during the middle jurassic regional uplift that affected the danish basin and the ringkøbing–fyn high, and thick paralic sandstones were deposited in the zone; these sandstones 18 fig. 2. generalised stratigraphic scheme of the danish onshore area along a nw–se-trending cross-section. the formations with potential aquifers are indicated in yellow and brown. note the pronounced erosion surfaces at the base of the middle jurassic and lower cretaceous and the progressive onlap to these surfaces. these features have a major influence on the regional distribution and burial depths of potential reservoirs. ag, stratigraphic position of the arnager grønsand formation; ngb, north german basin; rkf, ringkøbing–fyn high; skp, skagerrak platform; stz, sorgenfrei–tornquist zone. form excellent reservoirs (haldager sand formation; fig. 2). the zone is further characterised by pronounced late cretaceous – early tertiary tectonic inversion with the uplift of potential reservoirs. potential reservoirs the most promising reservoirs occur within the triassic – lower cretaceous succession (fig. 2). this succession has been the target of hydrocarbon exploration since 1935, and is thus known from about 60 deep wells and seismic data acquired over many years, although with a very variable data quality and coverage. based on regional geological studies (e.g. bertelsen 1978, 1980; michelsen et al. 2003; nielsen 2003) a number of stratigraphic units with a regional geothermal potential have been identified. these include the lower–upper triassic bunter sandstone and skagerrak formations, the upper triassic – lower jurassic gassum formation, the middle jurassic haldager sand formation and the upper jurassic – lower cretaceous frederikshavn formation. other formations may locally contain potential aquifers, such as the fine-grained sandstones of the f-ii member of the fjerritslev formation on the skagerrak– kattegat platform, and the arnager grønsand formation in easternmost zealand. the bunter sandstone formation is present south of the ringkøbing–fyn high, on parts of the high and in the danish basin. it grades into the skagerrak formation towards the north-eastern basin margin (bertelsen 1978, 1980). the bunter sandstone formation is dominated by fine-grained sandstones, mainly deposited in an arid continental environment dominated by fluvial channels, aeolian dunes and marginal marine facies. the skagerrak formation is less well known, but its marginal distribution along the northern and north-eastern basin margin, and the coarsegrained, often poorly sorted sandstones interbedded with claystones, suggest deposition in alluvial fans and lakes. the gassum formation is present in almost the entire danish area, and shows a remarkable lateral continuity with thickness generally between 100 and 150 m with a maximum of about 300 m in the sorgenfrei–tornquist zone (michelsen et al. 2003; nielsen 2003). the formation consists of fineto medium-grained, locally coarse-grained, sandstones interbedded with heteroliths, claystones and thin coals. the laterally continuous shoreface sandstones were deposited by repeated shoreline progradation. fluvial and estuarine sandstones dominate the lower–middle part of the formation in the sorgenfrei–tornquist zone. the haldager sand formation is up to 200 m thick in the sorgenfrei–tornquist zone, and shows a marked thinning towards the south-west and north-east (michelsen et al. 2003; nielsen 2003). it consists of thick, fineto coarsegrained sandstones alternating with thin siltstones, claystones and coals, deposited in shallow marine, estuarine, fluvial and lacustrine environments. the frederikshavn formation is present in the northern part of the danish area, and shows marked thickness variations (75–235 m), reaching a maximum in the sorgenfrei–tornquist zone (michelsen et al. 2003). the formation consists of siltstones and fine-grained sandstones interbedded with claystones. temperature and salinity of the formation water in these potential reservoirs increase with increasing depth. the temperature–depth relation is well established, and is rather uniformly developed over the danish area with a general gradient of about 30°c per km. the salinity shows a general increase of about 10% per km burial depth, but great variations are found. porosity and permeability decrease with increasing depth due to mechanical compaction and the formation of diagenetic minerals that reduce pore volume and pore connections. permeability is very critical, but difficult to predict since very large variations are found depending on depositional facies, provenance, mineralogical composition, burial history and position in the basin. these relationships and their mutual dependency are not fully understood, which weakens the predictive strength of the current geological models used for identifying areas of interest. however, combining the distribution of the above-described formations with an estimate of where sand thickness of the formations exceeds 25 m at depths of 1000–2500 m provides a useful indication of regional geothermal potential. figure 1 displays the potential for the land area of denmark in a general manner, and indicates which formations may warrant further investigation for geothermal energy production. existing and planned geothermal facilities the thisted plant in northern jutland is the only working geothermal plant in denmark, although a second plant is currently under construction in copenhagen (fig. 1). the thisted plant has produced heat from the gassum formation for almost 20 years without notable production or injection problems. a study of the geothermal potential in the copenhagen–malmö region was initiated in the year 2000 on behalf of dong a/s encouraged by financial support from the danish government and technological developments that make the utilisation of relatively low temperature formation water possible. the subsurface of the greater copenhagen area was previously poorly known, as no deep wells existed and seismic data coverage was very poor. new seismic data were acquired in 2001, and the survey has carried out a 19 20 geological evaluation of the geothermal potential at seven localities in the greater copenhagen area, based on integration of the new data with existing well and seismic data from denmark, øresund and southern sweden. the evaluation indicated the presence of several possible sandstone aquifers, including the gassum and bunter sandstone formations. the margretheholm location close to the centre of copenhagen was selected for further investigations, and a vertical well was drilled to about 2700 m in 2002 (fig. 3). the well encountered a promising aquifer in the bunter sandstone formation, and a second, deviated well was drilled to the same aquifer in 2003. the test results were promising, and a geothermal power plant is now under construction based on the utilisation of c. 70°c geothermal water. when established, the plant is expected to produce around 400 tj heat annually, corresponding to 1% of the total heating demand of the copenhagen area, with an option for future expansion. as a direct result of the successful efforts in copenhagen, geothermal exploration has now been resumed in other parts of the onshore danish area. geus is currently cooperating with dong a/s on the identification and assessment of several prospective sites. references bertelsen, f. 1978: the upper triassic – lower jurassic vinding and gassum formations of the norwegian–danish basin. danmarks geologiske undersøgelse serie b 3, 26 pp. bertelsen, f. 1980: lithostratigraphy and depositional history of the danish triassic. danmarks geologiske undersøgelse serie b 4, 59 pp. haenel, r. & staroste, e. (eds) 1988: atlas of geothermal resources in the european community, austria and switzerland. commission of the european communities, publication eur 11026, 74 pp., 110 plates. hurter, s. & haenel, r. (eds) 2002: atlas of geothermal resources in europe. european commission, publication eur 17811, 92 pp., 89 plates. michelsen, o. et al. 1981: kortlægning af potentielle geotermiske reservoirer i danmark. danmarks geologiske undersøgelse serie b 5, 28 pp. michelsen, o., nielsen, l.h., johannessen, p.n., andsbjerg, j. & surlyk, f. 2003: jurassic lithostratigraphy and stratigraphic development onshore and offshore denmark. in: ineson, j.r & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 147–216. nielsen, l.h. 2003: late triassic – jurassic development of the danish basin and the fennoscandian border zone, southern scandinavia. in: ineson, j.r & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 459–526. sørensen, k., nielsen, l.h., mathiesen, a. & springer, n. 1998: geotermi i danmark: geologi og ressourcer. danmarks og grønlands geologiske undersøgelse rapport 1998/123, 24 pp. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: lhn@geus.dk fig. 3. the drilling of two deep wells at margretheholm, central copenhagen. << /ascii85encodepages false /allowtransparency false /autopositionepsfiles true /autorotatepages /all /binding /left /calgrayprofile (dot gain 20%) /calrgbprofile (srgb iec61966-2.1) /calcmykprofile (u.s. web coated \050swop\051 v2) /srgbprofile (srgb iec61966-2.1) /cannotembedfontpolicy /warning /compatibilitylevel 1.4 /compressobjects /tags /compresspages true /convertimagestoindexed true /passthroughjpegimages true /createjdffile false /createjobticket false /defaultrenderingintent /default 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<feff4f7f752890194e9b8a2d5b9a5efa7acb76840020005000440046002065874ef65305542b8f039ad876845f7150cf89e367905ea6ff0c4fbf65bc63d066075217537054c18cea3002005000440046002065874ef653ef4ee54f7f75280020004100630072006f0062006100740020548c002000520065006100640065007200200035002e0030002053ca66f465b07248672c4f86958b555f3002> >> >> setdistillerparams << /hwresolution [2400 2400] /pagesize [595.000 842.000] >> setpagedevice geological survey of denmark and greenland. bulletin 10, 17-20 well-constrained depositional models are essential for successful exploration and field development. the skagen spitsystem offers a unique possibility for the establishment of a depositional model constrained by excellent outcrops, welldefined palaeogeography, good age control and detailed observations on hydrodynamics and morphology of the prograding part of the spit-system. the model offers a supplementary interpretation of shallow marine sandstones to the existing delta and linear shoreface models. the sand-dominated skagen spit-system is c. 22 km long, 4 km wide and up to 35 m thick, with a sand volume of c. 2.2 km3. if filled with oil, this system would contain 0.6 km3 corresponding to 3.8 x 109 barrels assuming a porosity of 30% and an oil saturation of 90%. this is comparable in size with the largest danish oil field (the dan field), in the north sea. reservoir models for isolated linear ‘offshore’ sandstone bodies have been controversial for many years. their size and internal indications of palaeocurrent directions are similar to those of the spit-system model, and this model may therefore be applicable for some of these bodies. depositional model for the skagen spit-system the skagen spit-system is part of the large triangular coastal complex termed skagen odde that forms the northern tip of jylland (fig. 1). the complex began to form at 7150 years bp, and the actual spit-system has developed within the past 5500 years (fig. 1); (nielsen & johannessen 2001, 2004). the older parts of the spit-system have been raised c. 13 m above present-day sea-level as a result of the rate of glacial rebound has exceeded the rate of eustatic sea-level rise during the holocene and can therefore be studied in cliff sections (fig. 2). the distal youngest part of the spit-system is still prograding by several metres per year, and the depositional processes can be studied at the point of the spit (fig. 3). age relationships are well constrained by c14 dates of peat and shells along the 22 km long spit-system (hauerbach 1992; clemmensen et al. 2001). the possibility of direct comparison between geological sections in the older raised part of the spit and the recent depositional processes at the point of the spit is unique, and provides a very high degree of certainty to the interpretation of the sedimentary units. the skagen spitspit-systems – an overlooked target in hydrocarbon exploration: the holocene to recent skagen odde, denmark peter n. johannessen and lars henrik nielsen © geus, 2006. geological survey of denmark and greenland bulletin 10, 17–20. available at: www.geus.dk/publications/bull 17 10 km 7150 bp l l l l l l l l l l l l l l l l l l l l l l l l l l l n n 10 km 50 m l l l l 5500 bp 4500 bp mainland pleistocene glacial till, fluvial sand, marine sand and mud skagen odde coastal elements pleistocene glacial till and fluvial sand late pleistocene coastal sand glacial till and fluvial sand covered by thin holocene coastal sand lagoonal sand and mud troldkær spit-system; sand and gravel strandplain; sand skagen spit-system; sand and gravel cliffs glacialdeposits holocene marine mud recent sand denmark fig. 1. geological development and palaeogeographic reconstruction of the skagen odde coastal complex. the skagen spit-system discussed here is shown in orange on the lowermost geomorphological map. the skagen spit-system began to form c. 5500 years bp. system may thus be regarded as a natural full-scale sedimentological laboratory. the spit-system overlies offshore mud. it shows a weak coarsening-upward trend and consists of three sand-dominated units locally topped by peat. the lowest consists of up to 25 m of storm sand beds (unit 1), and is overlain by c. 5 m of dune and bar-trough sediments (unit 2), followed by c. 2 m of beach sand and 0–2 m aeolian sand (unit 3), topped by peat lenses up to 1.5 m thick (unit 4; figs 2, 3). this spit-system succession is overlain by up to 10 m of recent aeolian sand. in the constructive phases of the development of the spit-system, sand is transported from eroding glacial deposits more than 40 km south-west of the present tip of the spit, 18 4: swale peat 3: beach sand 2: dune cross-beds 4 m 20 cm 1 m burrows of heart urchins storm sand bed 1: storm sand bed 40 cm fig. 3. photographs of the four depositional units and the sites of their formation (1–4) on the active subaerial spit and the submarine spit-platform at the point of the spit-system (section a–b). a b 2d & 3d dunes 2d & 3d dunes bar-trough system beach ridges storm sand current direction constructive wind directions shoreline of spit n a b storm sand beds beach sand with low angle slipface 0 0 20 m 1500 m beach ridge proximal bar with steep slipfaceswampy swale shoreline distal bar system 4 4 3 3 2 2 1 1 ne sw swale peat aeolian sand 2 m b ea ch sa n d d u n e & b ar -t ro u gh se d im en ts st o rm s an d b ed s 11 22 33 44 1 2 3 4 fig. 2. coastal cliff-section exposing the raised four depositional units of the skagen spit-system shown in fig. 3. and from the older uplifted part of the spit-system. the preferential wind directions are from the south-west and west creating a strong shore-parallel northward flowing current capable of transporting large amounts of sand (up to c. 1.5 million m3/year) along the coast. by far the greater part of the sediments are deposited in front of the spit point where storm sand beds are deposited from heavily loaded suspension currents in water depths of c. 7–30 m causing progradation (fig. 3). sandy 2d and 3d dunes migrate northwards along the spit coast, and pebbles are transported in the swash-backwash zone. deposition occurs mainly where the spit coast bends and refraction of the waves results in reduction of the transport capacity. at the same time, the longshore currents expand over the area at the tip of the spit, as the controlling effect of the subaerial spit ceases and water depth increases. this combination causes high sedimentation rates on the platform at the tip of the spit. the majority of the dunes migrate obliquely basinwards on the gentle seaward dipping platform surface in front of the spit, leaving behind thick units of cross-bedded sand that is deposited in water depths of c. 0.3–9.5 m (fig. 3). occasionally, a shore-attached bartrough system is formed in the surf zone along the front of the spit. such bars migrate towards the coast, emerge and become swash bars with swash-backwash lamination forming on the seaward side. during severe storms from the north, pebbly beach ridges are formed on the backshore up to c. 1 m above average sea-level. peat formation takes place in swales between beach ridges up to a few hundreds metres from the active spit coast. aeolian dunes subsequently develop and migrate across the spit. spit-systems may be differentiated from other shallow marine sandstones by the presence of platform foresets (e.g. nielsen et al. 1988), curved beach ridges showing pronounced lateral fining of grain-size, curved peat deposits, and palaeocurrent directions that differ from deltaic and linear shoreface sandstones (fig. 3). for instance, the foresets of the 2d and 3d dune sands in unit 2 show nearly unimodal palaeocurrent directions obliquely to the accretionary spit coast and therefore show palaeocurrent directions at high angles to the long, exposed side of the spit-system (figs 1, 3). preservation potential the subaerial part and the upper marine part of the spit-system may be subject to erosion during transgression, and parts of the aeolian sand, peat lenses and beach sand may be eroded away and replaced by a thin transgressive sand. however, the remaining part of the spit-system has a high preservation potential and may be enveloped and sealed by offshore mudstones. exploration and reservoir model for spit-systems headland-attached spit-systems linear coast and delta progradation systems depend on sediment input from rivers and distributary channels. spit-systems, on the other hand, rely on wave erosion and longshore drift and may be found downdrift from a headland (as the skagen spit-system) or a fault block partly submerged and exposed to waves. in the latter case spit-systems may form on a hanging-wall fault block down-drift from a footwall block (fig. 4). the thickness of the spit-system succession depends on the water depth in which it progrades; the platform will be thick above topographic lows and thin over highs. spitsystems have a tendency to prograde on top of elevated areas such as submarine ridges, because the progradation rate is greater at shallow water levels over the ridge than at deeper water depths on both sides of the ridge. consequently, spitsystems preferentially prograde on the plunging crest of fault blocks. a potential reservoir is therefore situated on the elevated crests rather than lying in the deep parts of the hanging-wall blocks (fig. 5). detached spit-systems – a model for offshore sand bars? during progradation of a spit-system the headland and the proximal part of the spit may be eroded by wave activity and 19 spit-system fig. 4. spit-system formation and preservation on a hanging-wall fault block. fig. 5. conceptual model of spit-systems attached to plunging fault block crests and prograding on the submarine part of the fault block crest. longshore currents, and the spit-system may eventually be detached and become an isolated sand body (fig. 6). the upper part of the sand body will be exposed to erosion, and only the lower part of the former spit-system may be preserved. elongated, isolated sandstone bars (up to 35 km long, 2–4 km wide and 30 m thick) encased in offshore mudstones and apparently deposited some distance from land are common hydrocarbon exploration targets in the western interior seaway of the usa (e.g. suter & clifton 1999). however, their genesis has been controversial for decades, and the reservoir models are poorly constrained. some of these sandstone bodies contain unimodal cross-bedded units which indicate palaeocurrents at an acute angle to the length of the bars. their size, facies and palaeocurrent directions are similar to those of the spit-system model. concluding remarks the middle and upper jurassic in the north sea rift basins are characterised by heavily block-faulted areas subjected to transgressions. as extensive spit-systems can develop within a few thousand years under the right conditions, it is likely that spit-systems were locally formed on partly submerged, fault block crests and on hanging-wall fault blocks. aspects of the skagen spit-system model have been applied in the mapping of reservoirs in the troll field in the norwegian north sea (dreyer et al. 2005). acknowledgements norsk hydro and the carlsberg research foundation (ans-0881/10) provided financial support to the studies of the skagen spit-system. references clemmensen, l.b., richardt, n. & andersen, c. 2001: holocene sea-level variation and spit development: data from skagen odde, denmark. the holocene 11, 323–331. dreyer, t., whitaker, m., dexter, j., flesche, h. & larsen, e. 2005: from spit systems to tide-dominated delta: integrated reservoir model of the upper jurassic sognefjord formation on the troll west field. in: doré, a.g. & vining, b. (eds): petroleum geology: north-west europe and global perspectives. proceedings of the 6th petroleum geology conference, 423–448. london: geological society. hauerbach, p. 1992: skagen odde – skaw spit. an area of land created between two seas. folia geographica danica 20, 119 pp. nielsen, l.h. & johannessen, p.n. 2001: accretionary, forced regressive shoreface sands of the holocene–recent skagen odde spit complex, denmark – a possible analogue to fault-attached shoreface sandstone reservoirs. in: martinsen, o. & dreyer, t. (eds): sedimentary environments offshore norway – palaeozoic to recent. norwegian petroleum society (npf) special publications 10, 457–472. nielsen, l.h. & johannessen, p.n. 2004: skagen odde – et fuldskala, naturligt laboratorium. geologi – nyt fra geus 1, 1–12. nielsen, l.h., johannessen, p.n. & surlyk, f. 1988: a late pleistocene coarse-grained spit-platform sequence in northern jylland, denmark. sedimentology 35, 915–937. suter, j.r. & clifton, h.e. 1999: the shannon sandstone and isolated linear sand bodies: interpretations and realizations. in: bergman, k.m. & snedden, j.w. (eds): isolated shallow marine sand bodies: sequence stratigraphic analysis and sedimentologic interpretation. sepm (society for sedimentary geology) special publication 64, 321–356. 20 author’s address geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: pjo@geus.dk sl-1 sl-1 sl-2 sl-1 a a' b b' offshore mudstone offshore mudstone transgressive sand cover isolated sub-marine sand body sea level time 1 ravinement surface sedimentary spit units active spit-system a'a b'b lowstand spit-system attached to mainland transgression drowned, detached, ravined spit-system encased in marine mudstones apparent distance to coeval shoreline obliquely in-comming waves strong shore-parallel current 2 3 4 1 1 41 2 _ fig. 6. conceptual model of detached spit-systems. isolated, elongated sandstone bars in the western interior seaway (usa) may originally have been spit-systems prograding from a headland (a–a’). later they became detached from the headland due to erosion of the headland and the proximal part of the spit-system. during subsequent transgression the headland became submerged and upper parts of the spit-system was eroded (b–b’) and the preserved sandstone bars may appear to have formed at great distance from the mainland. geological survey of denmark and greenland bulletin 1, 247-264 247 middle jurassic – early cretaceous rifting of the danish central graben jens j. møller and erik s. rasmussen during the jurassic – early cretaceous, the danish central graben developed as a n–sto nnw– sse-trending graben bounded by the ringkøbing–fyn high towards the east and the mid north sea high towards the west. the graben consists of a system of half-grabens and evolved by faultcontrolled subsidence; three main rift pulses have been recognised. the first pulse ranged from the callovian to the early oxfordian, the second pulse was initiated in the latest late kimmeridgian and lasted for most of the early volgian, and the third and final pulse occurred within the ryazanian in the early cretaceous. the first pulse was characterised by subsidence along n–s-trending faults. the most pronounced fault-controlled subsidence occurred in the east, especially along n–s-striking segments of the boundary fault to the ringkøbing–fyn high. during this period, minor salt movements occurred with the development of salt pillows. the activity along the n–s-trending faults ceased during the oxfordian. during the second pulse, in early volgian times, subsidence was concentrated along new nnw–sse-trending faults and the main depocentre shifted westward, being most marked within the tail end graben, the arne–elin graben, and the feda graben. this tectonic event was accompanied by the accumulation of a relatively thick sediment load resulting in the development of salt diapirs, especially within the salt dome province. the third tectonic pulse was essentially a reactivation of the nnw–sse-trending structures and there is clear evidence of subsidence controlled by faulting and salt movements. despite the overall extensional tectonic regime, local compressional tectonics resulted in thrusting. for instance, the gert ridge is interpreted to have formed by readjustment at the boundary fault between two subsiding blocks. the structural framework during graben evolution controlled, to some degree, the distribution of reservoir sandstones. reservoir sandstones associated with periods of rotational tilt include middle jurassic deposits referred to the bryne and lulu formations, and upper jurassic sandstones referred informally to the ‘fife sandstone formation’. sands deposited during tectonic relaxation are represented by the heno formation and upper jurassic turbidites interbedded in the farsund formation. sea-level changes were probably most important during periods of tectonic relaxation, particularly with respect to the deposition of lowstand sandstones in basinal areas. keywords: north sea, danish central graben, middle jurassic – lowermost cretaceous, 3d seismic data, structural evolution, reservoir sandstone distribution geological survey of denmark and greenland, geocenter copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: jjm@geus.dk the structural development of the danish part of the central graben during the jurassic has previously been described by andersen et al. (1982), gowers & sæbøe (1985), møller (1986), cartwright (1987), and korstgård et al. (1993). the structural evolution of the norwegian central trough (graben) has been presented by gowers geological survey of denmark and greenland bulletin 1, 247–264 (2003) © geus, 2003 248 et al. (1993). these papers are based on interpretation of 2d multichannel seismic data acquired during the early 1980s. despite the relatively high quality of these data, new information has come to light as a result of recentlyacquired 3d seismic data and information from wells drilled in the late 1980s and in the early part of this decade. the purpose of this paper is to present our interpretation of the evolution of the danish central graben, on the basis of these new data. furthermore, mapping of oil and gas fields demonstrates a close link between the structural framework and the location of reservoir rocks, in this case, sandstones. therefore, the influence of the structural framework on the distribution of reservoir sandstones is discussed. geological framework structural development the danish central graben is a major east-dipping halfgraben system bounded towards the east by the ringkøbing–fyn high and towards the west by the mid north sea high (fig. 1). the half-graben consists of a series of minor nw–se-trending half-grabens mainly dipping eastwards. the graben formed between the middle jurassic and the earliest cretaceous by the interaction of basement faulting and mobilisation of zechstein and triassic(?) salt (gowers & sæbøe 1985; møller 1986; cartwright 1987; korstgård et al. 1993). locally, stretching factors have been calculated to be up to two (vejbæk 1992). the faulting and salt mobilisation resulted in a complex configuration of highs and graben features, the most distinct of which have been named (andersen et al. 1982; møller 1986; vejbæk 1986; japsen et al. 2003, this volume). tectonic movements in the central graben area have been recognised to have occurred as early as the permian (ziegler 1975; gowers & sæbøe 1985); indeed, rotliegendes strata in the danish sector record rifting and associated volcanism although not restricted to the central graben area (stemmerik et al. 2000). zechstein deposits are present in the northern and southern parts of the danish central graben (e.g. the søgne basin and the salt dome province) but are probably absent in the central part. there may have been a connection between the southern and northern permian basins through the tail end graben during zechstein times although evidence (e.g. salt structures) is lacking. triassic sediments were more widespread and relatively thick triassic deposits were probably laid down, as indicated by the u-1 well (jacobsen 1982) and discussed by vejbæk (1992). tectonic quiescence prevailed during the early jurassic and a uniform succession of marine mudstones accumulated (michelsen 1978; andsbjerg & dybkjær 2003, this volume). middle jurassic updoming of the central graben area (ziegler 1982) resulted in significant erosion of triassic and lower jurassic sediments. following middle jurassic uplift, a stretching phase caused rapid fault-controlled subsidence, particularly in the tail end graben area (vejbæk 1992). this mainly late jurassic rifting ceased during the early cretaceous and it has been proposed that subsequent subsidence was controlled by a combination of thermal relaxation and gabbro–eclogite phase transformation, the latter process explaining the increased subsidence rate during the cenozoic (vejbæk 1992). kooi et al. (1991), on the other hand, suggested that increased compressional stresses within the north sea region were the main cause of accelerated subsidence in the late cenozoic. structural elements the complex pattern of the danish central graben has made it necessary to name the most important structural elements (fig. 1); the following usage is based on britze et al. (1995) and japsen et al. (2003, this volume). towards the east, the danish central graben is bounded by a segment of the ringkøbing–fyn high termed the east north sea high. the eastern part of the graben is subdivided into the søgne basin towards the north, the tail end graben in the central part, and the salt dome province and the rosa basin in the south. minor elements such as the poul plateau constitute the transition between the main boundary fault and the graben itself. towards the west, the tail end graben passes into the heno plateau. in the northern part of the graben, a more complex development occurred involving the evolution of a series of grabens, the arne–elin, gertrud and feda grabens. the søgne basin is separated from the gertrud graben by the mandal high and the piggvar terrace. the mid north sea high forms the eastern limit of the danish central graben; at the transition this is segmented into the mads and inge highs. stratigraphic evolution of the danish central graben pre-middle jurassic the p-1 well located on the mads high terminated in caledonian basement of greenschists. the oldest sedimentary deposits known in the danish central graben are of carboniferous age, as encountered in the p-1 well (michelsen 1982) and in the gert-2 and gert-3 wells. thick carboniferous sections are known both south and north of danish territory (michelsen 1982) and the carboniferous is believed to have been distributed over most of the area, but was probably eroded during later tectonic events. the lower permian section is characterised by volcanic rocks succeeded by aeolian deposits and sabkha sediments of the rotliegendes group (jacobsen & larsen 1982; stemmerik et al. 2000). transgression of the northern and southern parts of the danish central graben in the late permian resulted in the deposition of carbonates and evaporites referred to the zechstein group. triassic sedimentation occurred in basinal areas inherited from the permian, but elevated areas were also periodically areas of deposition. the triassic section is dominated by non-marine sandstones 249 20 km reverse fault normal fault salt diapir salt pillow well fault plane n dk gnl uk n 500 km 4°e 56°n fig. 5 fig. 13 fig. 4 fig. 7 fig. 3 søgne basin gertrudgraben feda graben inge high outer roughbasin heno plateau m ads h igh poulplateau salt dome province coffee soil fault igor basin r osa basin a rne–elin g raben piggvar terrace ringkøbing–fyn h igh east n orth sea h igh ål basin tail end graben mid north sea high n. jens-1 u-1 nora-1 gwen-2 cleo-1 gert-4,-1,-3,-2 jeppe-1 p-1 tordenskjold-1 ravn-2 ravn-1 elly-1,-2 gert ridge mandal high fig. 1. location map and structural elements of the danish central graben. wells referred to in the text are indicated, together with the location of the geosections in figs 3–5 and the seismic lines in figs 7, 13. position of additional wells named on figs 3–5 is also shown. national sectors of the north sea: dk, denmark; g, germany; n, norway; nl, the netherlands; uk, united kingdom. 250 and shales with subordinate evaporites (jacobsen 1982) and in some areas is up to 2000 m thick. a relatively uniform succession of marine shales is believed to have been deposited over most of the danish north sea region during the early jurassic (koch et al. 1982; michelsen et al. 2003, this volume), but most of this was eroded in the middle jurassic. middle–late jurassic the collapse of the north sea plume (ziegler 1982; underhill & partington 1993) and graben formation resulted in progressive marine transgression of the area in the middle and earliest late jurassic. deposition of fluvial, lacustrine, and nearshore sandstones occurred system stage lo w er ju ra ss ic u pp er m i d d l e l o w e r åsgard formation leek member bo member heno fm lola formation middle graben formation bryne formation lithostratigraphy ryazanian volgian kimmeridgian oxfordian callovian bathonian bajocian aalenian toarcian pliensbachian sinemurian hettangian c re ta ce ou s valanginian fjerritslev formation vyl fm poul fmfarsund formation lulu formation structural evolution subsidence along nnw–sse faults subsidence along nnw–sse faults fault-controlled subsidence along n–s faults u l u l l u u m l l m u m l u l u u l l u m l u u l u m l hiatus fig. 2. jurassic – lowermost cretaceous lithostratigraphy of the danish central graben (modified from michelsen et al. 2003, this volume). offshore organic-rich marine shales marine mudstones and siltstones submarine fan sandstones and siltstones shallow marine sandstones and siltstones paralic and non-marine sandstones, siltstones, mudstones and coals marine calcareous mudstones and marlstones 251 during this period, such as the ‘basal sandstone unit’ in the norwegian sector of the feda graben (bergan et al. 1989), the bryne formation in the salt dome province and the tail end graben, and the bryne and lulu formations in the søgne basin (damtoft et al. 1992; johannessen & andsbjerg 1993; michelsen et al. 2003, this volume; fig. 2). these sand-dominated formations are succeeded by marine shales of the lola formation and sandstones of the heno formation. the latter formation has an overall retrogradational stacking pattern although widespread progradation of sands occurred during the late kimmeridgian, represented by the upper part of the heno formation (the ravn member, see johannessen 2003, this volume; michelsen et al. 2003, this volume). following regional flooding in the latest late kimmeridgian, a thick mudstone-dominated marine succession (the farsund formation) was deposited. the farsund formation ranges in age from latest late kimmeridgian to early ryazanian and is very rich in organic matter at a number of levels; it forms the most important hydrocarbon source rock in the danish area. of particular note is the bo member within the uppermost part of the farsund formation which is an excellent source rock (ineson et al. 2003, this volume). turbiditic sandstones have been recognised in several wells in the upper farsund formation, at roughly the same stratigraphic level as the organic-rich mudstones of the bo member (damtoft et al. 1992; andsbjerg & dybkjær 2003, this volume). data and methods the present study is based on all multichannel 2d seismic surveys acquired since 1980 (rtd, sp, dmk, dk, cgd, np, nh), giving a very dense coverage of the study area. in addition, selected 3d surveys have been utilised in the study. well information from 43 exploration wells penetrating the jurassic succession have been used. well-to-seismic ties have been achieved using the velocity surveys from released wells (nielsen & japsen 1991). the age assignments of the sedimentary succession are based on recent studies (johannessen et al. 1996; andsbjerg & dybkjær 2003, this volume). the seismic interpretation is a combination of traditional mapping of distinct reflectors or unconformities and seismic facies analysis. regionally mappable reflectors have been correlated with the sequence stratigraphic subdivision of andsbjerg & dybkjær (2003, this volume). seismic facies analysis has been applied to selected seismic sections, incorporating recent idealised models of seismic facies characteristics with respect to the recognition of tectonic systems tracts (e.g. prosser 1993). the systems tracts have then been correlated to wells, wherever possible, although such correlations often involve major uncertainties since wells are normally located on structural highs where seismic ties are poor. the most important wells used in this study were the gert-4, nora-1, and north jens-1 wells. structural evolution to illustrate the overall structural framework of the danish central graben, three geosections are presented covering the southern, middle and northern part of the graben, together with an isochore map of the upper jurassic succession (figs 3–6); these illustrations are based primarily on older 2d seismic surveys integrated with borehole data. it must be stressed, however, that the structural interpretation presented here is based on new high resolution seismic data, such as that shown in figure 7 and the study of the 3d seismic data from the gert area. consequently, there are some discrepancies between the seismic units displayed on the geosections and the rift events proposed in this paper. in general, accumulation of the thickest jurassic succession was associated with marked subsidence along the coffee soil fault. this subsidence pattern was modified by salt movements during the jurassic, especially in the southern part, the so-called salt dome province. a relatively thick jurassic interval is also seen in the gertrud and feda grabens (figs 5, 6). this present configuration of the graben fill is the result of successive fault activity resulting in combined fault-controlled subsidence of half-grabens and subsidence/uplift related to salt movements. from the middle jurassic to the earliest early cretaceous, three tectonic pulses are recognised: (1) callovian to early oxfordian fault-controlled subsidence, concentrated along n–s-striking faults; (2) a latest late kimmeridgian to early volgian pulse, characterised by a distinct shift to subsidence along nnw–sse-trending faults; (3) a final pulse (ryazanian) also involving fault-controlled subsidence along nnw–sse-striking faults. tectonic pulse 1 the first pulse, from the callovian to the early oxfordian, was characterised by mainly fault-controlled subsidence along n–s-striking faults. such faults include segments ravn-1 elin-1 nora-1 heno plateau tail end graben coffee soil fault 1 0 2 3 4 5 6 7 8 9 10 20 30 d ep th ( km ) 40 50 (km) wsw ene fig. 4. geosection of the tail end graben and the heno plateau (for location, see fig. 1; for legend, see fig. 3). kim-1 lone-1 rita-1 gert-2 gert-3 mona-1 karl-1 w. lulu-4 w. lulu-2 lulu-1 cleo-1 elna-1 mid north sea high feda graben gertrud graben mandal high søgne basin coffee soil fault ringkøbing–fyn high (east north sea high) 1 0 2 3 4 5 6 7 8 10 20 30 40 d ep th ( km ) 50 60 70 (km) wsw ene gert ridge fig. 5: geosection of the northern part of the danish central graben (for location, see fig. 1; for legend, see fig. 3). the gert-3 well lies just north of the section and is projected into the line of section. john-1 john flank-1 nils-1a-2 o-1 post-ryazanian upper volgian – ryazanian lower–upper volgian upper kimmeridgian – lower volgian callovian – upper kimmeridgian aalenian–callovian triassic – lower jurassic zechstein pre-zechstein coffee soil fault 1 2 3 4 5 6 7 1 0 2 3 4 5 7 8 10 20 30 40 50 (km) d ep th ( km ) ew fig. 3. geosection of the salt dome province (for location, see fig. 1). of the coffee soil fault, n–s-striking faults within the tail end graben (figs 7, 8), and faults in the north-western area around the gert ridge (rasmussen 1995). in the tail end graben (fig. 7), the seismic reflection pattern records the response of the sedimentary system to different stages in rift evolution; two discrete tectonic events can be identified in this area (tectonic pulses 1, 2). the base of the first event is characterised by a distinct angular unconformity, created by erosion of the footwall (fig. 7). above the unconformity, two wedgeshaped reflection packages are ascribed to the first tectonic pulse; they are interpreted to reflect different stages in the rotation of the fault block. as illustrated in the flattened section (fig. 9), the lower wedge-shaped interval (rotational tilt package 1) is characterised by a low amplitude and discontinuous reflection pattern, which grades upwards into a high amplitude, continuous and wedge-shaped dipping reflection pattern that can be followed over the entire tilted fault block area. the upper wedge (relaxation package 1) is characterised by a low amplitude, gently-dipping parallel and sub-parallel seismic reflection pattern, showing regular onlap onto the underlying interval. little indication 253 5°e 56°n 55°30'n 4°e 25 km upper jurassic isochore map contour interval 400 m 1:880 000 utm zone 31 normal fault at ‘base upper jurassic’ level reverse fault at ‘base upper jurassic’ level fault plane jurassic thin or absent well reaching base upper jurassic thickness in metres 0 2000 4000 fig. 6: isochore map of the upper jurassic succession of the danish central graben (from japsen et al. 2003, this volume). 254 of thickening of individual units can be recognised within this package. the seismic reflection pattern can thus be interpreted to reflect episodic block rotation during rifting. the lower wedge-shaped reflection pattern, showing thickening of individual stratal units towards the down-tilted area reflects deposition during active block rotation (prosser 1993, fig. 8). the overlying wedge, showing a parallel to sub-horizontal reflection pattern, is interpreted to represent passive fill during tectonic relaxation or a period during which the rate of block rotation was strongly reduced and the sediment mainly filled the topographic low created during the main rift pulse. the prominent onlap of the relaxation package onto a dipping surface, however, indicates that major tilting must have occurred prior to deposition of the sediments. it is therefore possible that the rift climax was characterised by non-deposition or that syn-rift sediments were confined only to the lows, where seismic resolution is poor in this area. measurements of recent tilting rates of fault blocks in greece (gawthorpe et al. 1994) show that the process of active tilting can be very rapid, up to 5 mm/year = 5000 m/ma. deposition of the jurassic – lowermost cretaceous succession described above occurred over a period of 29 ma; given a total thickness of c. 4000 m, the average tilt corresponds to 137 m/ma, which is significantly less than that suggested by recent measurements. it is thought likely, n –s -t re nd in g fa ul t n w –s etr en di ng fa ul t ? 500 ms 2 km sw ne intra-ryazanian intra-volgian latest late kimmeridgian early oxfordian late bathonian – callovian near-base middle jurassic faults post-rift package relaxation package 2 rotational tilt package 2 relaxation package 1 rotational tilt package 1 pre-rift package fig. 7: a ne–sw-striking seismic line through the tail end graben (for location, see fig. 1). the first two rift pulses described in the text are evident (1, 2), each being divisible into two packages: (a) a lower package that records active rotational tilting and is characterised by wedgeshaped stratal units that thin away from the footwall and (b) an upper ‘relaxation’ package that shows onlap onto the dipping surface of the rotated block. note that the first rotational tilt package (1) truncates the underlying pre-rift jurassic strata and that the n–s-striking fault became inactive in the early oxfordian during the relaxation phase (relaxation package 1). note also that the third pulse of block rotation cannot be demonstrated in this area. 255 therefore, that the main phase of tilting was so fast that sedimentation never managed to fill the space created. consequently, rift-climax deposits are volumetrically insignificant and not resolvable on seismic data. the faults of the first rift pulse were active during the deposition of the bryne and middle graben formations in the salt dome province and the bryne and lulu formations in the tail end graben and søgne basin (j. andsbjerg, personal communication 1998). the relaxation package is represented by the lola formation in the above-mentioned depocentres. the lola formation was also deposited on the heno plateau where it is succeeded by the sand-dominated heno formation. in the feda graben, the gert member (heno formation), the lola formation and the ravn member (heno formation) together represent the early graben fill corresponding to deposition during tectonic quiescence (johannessen et al. 1996). biostratigraphic data from wells that encountered the sediments deposited during the first tectonic pulse suggest that the faults were probably active from ?middle jurassic times (andsbjerg & dybkjær 2003, this volume) 55°30' 56°00' 5°30'5°00'4°30'4°00'3°30' 0 500 1000 1500 2000 well in which the succession was penetrated well in which the succession was not penetrated/encountered salt diapir fault trend fault plane callovian – upper kimmeridgian thickness (m) 25 km fig. 8. isochore map of the callovian – upper kimmeridgian succession showing the main faults that were active in this period. note firstly the influence of n–s-striking faults, which were active during the callovian and early oxfordian, and secondly that depocentres are close to the faults. 256 although the age and duration of the fault activity is difficult to determine accurately. however, in the tail end graben the unconformity illustrated in figure 7 probably truncates middle jurassic lacustrine sediments of the bryne formation and is overlain by rotational tilt deposits of the uppermost bryne and lulu formations indicating onset of rifting in the late bathonian – callovian (middle jurassic) and cessation in the early oxfordian (late jurassic). the tectonic pattern characterised by the n–s-striking faults probably represents reactivation of older faults. most of the half-grabens formed during this phase were tilted towards the east, although the feda graben formed as a westwards-dipping half-graben bounded by the mid north sea high towards the west. the gertrud graben was a later feature (see below) and at this time formed a plateau region that was continuous with the heno plateau separating the feda graben from the tail end graben. although acting as a plateau, minor faulting probably accompanied the major fault activity in the graben structures. tectonic pulse 2 following a period of tectonic quiescence, discrete subsidence occurred along nnw–sse-trending faults (figs 7, 10) during the latest late kimmeridgian and early volgian. this resulted in displacement along the coffee soil fault from the salt dome province in the south and further northwards. however, the northernmost segment of the coffee soil fault, which previously bounded the søgne basin, became inactive and displacement occurred along the nw–se-striking fault just west of the søgne basin (fig. 10). increased thicknesses within the søgne basin, as indicated on the late kimmeridgian – early 500 ms 2 km sw ne relaxation package 1 rotational tilt package 1 pre-rift package latest late kimmeridgian fig. 9. the central portion of the seismic section shown in fig. 7, flattened on the boundary between the first two rift pulses (late late kimmeridgian surface). the figure thus shows the stratal pattern of the first rift pulse. for key to reflectors, see fig. 7. 257 volgian isochore map (fig. 10), are documented by the cleo-1 well which penetrated a lower volgian succession. however, seismic data reveal that this volgian succession represents a tectonic relaxation package that was preserved at the cleo-1 well location as a result of later tilting of the søgne basin (during tectonic pulse 3). during the second pulse, fault-controlled subsidence occurred in the arne–elin graben (fig. 11) and along the faults bounding the feda and gertrud grabens (fig. 5). the occurrence of fault-controlled segmentation of the heno plateau is demonstrated by the evidence of increased subsidence at the site of the ravn-2 well compared with the nearby ravn-1 site (fig. 11; andsbjerg & dybkjær 2003, this volume). segmentation of the mid north sea high was also initiated at this time, resulting in the formation of the mads and inge highs and the basins west of the mads and inge highs (mackertich 1996). in the tail end graben, the depocentre shifted basinwards reflecting the influence of salt withdrawal concurrent with fault activity. during the tectonic quiescence that followed the second tectonic pulse, the relief created was passively infilled by the upper farsund 55°30' 56°00' 5°30'5°00'4°30'4°00'3°30' well in which the succession was penetrated well in which the succession was not penetrated/encountered salt diapir fault trend fault plane upper kimmeridgian – lower volgian cleo-1 12001000 14008006004002000 25 km thickness (m) fig. 10. isochore map of the upper kimmeridgian – lower volgian succession. note that the depocentres have shifted to occur along nw–se-trending faults in the tail end graben and gertrud graben (cf. fig. 8). the depocentres still lie close to the faults. 258 formation, including the organic-rich mudstones of the bo member (fig. 12). tectonic pulse 3 the third and final tectonic pulse in the formation of the danish central graben occurred in the midryazanian (early cretaceous). rotation of fault blocks and deep erosion on footwall crests resulted in the formation of a distinct angular unconformity which is best seen in the northern part of the graben where the influence of salt movements was insignificant (fig. 13). the age of this event is poorly constrained, but upper volgian – lower ryazanian organic-rich sediments of the upper farsund formation (bo member) are rotated on the heno plateau (ravn-2) and in the gertrud graben (gwen-2 and jeppe-1 wells) suggesting that the tectonic pulse post-dates these lower ryazanian deposits. the topography formed during this pulse was filled by an onlapping lower cretaceous succession (cromer knoll group). the lowermost sediments of this group were deposited in latest ryazanian – early valanginian times in this area (vejbæk 1986), thus suggesting that this 55°30' 56°00' 5°30'5°00'4°30'4°00'3°30' well in which the succession was penetrated well in which the succession was not penetrated/encountered salt diapir fault trend fault plane middle – lower upper volgian thickness (m) 10008006004002000 ravn-1 ravn-2 25 km fig. 11. isochore map of the middle and lower upper volgian succession. note the displacement of the depocentres away from the master fault. 259 third tectonic pulse occurred in the mid-ryazanian. salt tectonics were involved in this phase, especially in the feda graben where salt movements began in the volgian. the basinwards shift of depocentres in the tail end graben also reflects involvement of salt in this part of the danish central graben (fig. 12). the extensional tectonic pattern during this phase, however, resulted in local compression between different blocks; this effect was most marked between the two opposite-dipping feda and gertrud grabens where the gert ridge was formed (rasmussen 1995). it is uncertain whether this local thrusting was the result of strike-slip movements or local compensation between two subsiding blocks. strike-slip movements have been suggested for the formation of the arne–elin graben (clausen et al. 1996). the existence of local thrusting in the late jurassic has also been suggested by gowers et al. (1993) for the norwegian central graben. regional graben evolution the three-fold evolution of the danish central graben described above is comparable in many respects to the well in which the succession was penetrated well in which the succession was not penetrated/encountered salt diapir fault trend 55°30' 56°00' 5°30'5°00'4°30'4°00'3°30' 100 200 300 400 5000 uppermost volgian – ryazanian thickness (m) 25 km fig. 12. isochore map of the uppermost volgian – ryazanian succession. 260 development of the southern norwegian sector of the central graben, as described by gowers et al. (1993). the similarities involve not only the timing of tectonic pulses but also the fault trends along which active subsidence occurred. however, the earliest cretaceous (mid-ryazanian) event, involving tilting of minor fault blocks within former half-grabens, is probably better constrained in areas with insignificant salt tectonics, such as the northern part of the danish area. discussion structural evolution and sedimentation the structural evolution of a basin results both in the creation of accommodation space for the sedimentary fill and in the elevation of potential sedimentary source areas. the development of the sedimentary succession with respect to different tectonic pulses or phases in graben systems has been described by many workers, including surlyk (1989), prosser (1993), nøttvedt et al. (1995), ravnaas et al. (1997), rasmussen et al. (1998) and ravnaas & steel (1998). these papers deal with the sedimentary stacking patterns that result from the different stages in the rift evolution focusing on changes in accommodation space and variation in sediment supply. in addition, sequence stratigraphic analysis may provide a tool for stratigraphic and spatial lithological prediction within a rift basin based on the response of the system to relative sea-level changes. however, in a complex rift basin such as the danish central graben, which evolved through three tectonic pulses along different fault trends, prediction of the distribution of potential reservoir facies within the basin is difficult and an understanding of the structural framework is of critical importance. 2 km 500 ms wsw ene intra-ryazanian latest late volgian intra-volgian latest late kimmeridgian late kimmeridgian faults fig. 13: seismic section from the gertrud graben (for location, see fig. 1) showing the result of intense tilting of fault blocks during the third and final tectonic pulse in the mid-ryazanian. note the prominent onlap at the base of the cromer knoll group (?upper ryazanian – albian) indicating deposition after rotation of fault blocks. the period of fault block rotation was thus short-lived as indicated by onlap of ?upper ryazanian sediments in topographic lows (towards the ne). deposition during active block rotation the rotational package of the first rift pulse (fig. 7) is represented by the fluvial deposits of the uppermost bryne and lulu formations in the northern part of the danish central graben and by the uppermost bryne formation and the middle graben formation in the southern part. during deposition of these formations, the drainage systems were transverse (andsbjerg 2003, this volume) such that fluvial sands were concentrated in the elevated areas and lacustrine or marine finegrained sediments accumulated in the deeper parts on the hangingwall blocks. it is therefore not anticipated that significant fluvial sands were deposited adjacent to the master fault, as suggested for axial-trending fluvial sandstones (alexander & leeder 1987). on the mid north sea high, in the westernmost part of the danish territory, only one released well was available at the time of writing, namely the tordenskjold-1 well. this well did not encounter reservoir sands in the jurassic succession, but in the british sector, immediately west of the danish area, shallow marine reservoir sandstones of relatively high quality are known in the fife field (mackertich 1996). these sandstones, informally named the ‘fife sandstone formation’, were deposited on a slowly subsiding shelf during early to middle volgian times thus representing the rotational tilt package of the second rift pulse. the thickness variation of this formation was primary controlled by nnw–sse-trending faults, and well data show sandstone thickening and improvement in reservoir quality towards the main fault. on the footwall crest, a much thinner fife sandstone succession has been encountered. this sandstone was deposited during the second tectonic pulse and the thickest and best-developed sandstones were laid down on the hangingwall of the fault. deposition during tectonic quiescence in the gert and elly fields, the reservoir sandstones belong to an overall onlapping succession deposited during tectonic quiescence after the first rift pulse (fig. 7; see also fig. 13 in johannessen et al. 1996). these sandstones were deposited in nearshore depositional environments and unconformably overlie either middle jurassic continental sediments or older rocks. coastal sands were therefore successively deposited on elevated areas in the basin and at the basin margin. in the gert area, reservoir-quality sands were deposited in the late kimmeridgian adjacent to n–s-striking faults related to the first rift pulse. these faults controlled the distribution of the sand during deposition suggesting that differential subsidence across certain n–s faults persisted into the kimmeridgian despite becoming regionally inactive in the oxfordian. younger nw–se-trending faults subsequently created the gert ridge (søderstrøm et al. 1991; sundsbø & megson 1993; rasmussen 1995) and defined the structural closure that hosts the gert field. appraisal wells drilled on the gert structure have to date been located on the structural high and the thickness of reservoir section encountered has been disappointing. the influence of the n–s-striking fault should be incorporated in future field development models and may help in predicting the distribution of these reservoir sandstones. the gert area, furthermore, demonstrates the close relationship between a local sediment source area and deposition of a thick reservoir section. uplifted carboniferous fluvial sandstones in the footwall probably formed the source for the gert member sandstones in this area. this is suggested by the presence of carboniferous sandstones beneath the upper jurassic succession in the gert-2 well and the abundance of reworked carboniferous palynomorphs in the gert member in the gert-1 well (k. dybkjær, personal communication 1997). another interesting aspect related to the development of the lola and heno formations is that due to the overall relative sea-level rise during the late oxfordian and kimmeridgian, progradation during tectonic quiescence was muted; the progradational pulse observed in the uppermost part of the heno formation was probably related to sea-level stillstand or fall rather than to the tectonic evolution of the basin. the second tectonic pulse resulted, as for the first pulse, in the formation of a basin topography that was passively filled during the following period of tectonic quiescence. however, in this case the danish central graben was fully submerged and marine conditions prevailed, with the exception of marginal areas of the ringkøbing–fyn high and the mid north sea high. the upper volgian – ryazanian turbidites encountered in the jeppe-1 and iris-1 wells represent a relatively untested play within the danish central graben. the turbidites have been interpreted as passive fill succeeding the second tectonic pulse (rasmussen et al. 1999). as these turbidites were deposited in a basin with marked relief, it is anticipated that the sands were concentrated in the axial parts of the subbasins, as seen for example in the pendle grit of the british namurian (collinson 1988). the third and final tectonic pulse resulted in block rotation 261 and marked erosion on footwall crests. in the northern and western part of the danish central graben, the sand-rich heno formation, the lower–middle volgian nearshore sandstones on the mid north sea high (mackertich 1996) and upper volgian – ryazanian turbidites were eroded. reworked coarse-grained sediments may therefore have been deposited within lows formed in association with this tectonic pulse and may represent a lower cretaceous sandstone play that so far has not been tested in the danish central graben. concluding remarks it appears therefore that the tectonic framework is very important for the distribution and quality of reservoir rocks. under terrestrial conditions, in the middle jurassic, the reservoir sands accumulated on the upper and middle parts of the hangingwall slope. in the lower part of the hangingwall slope and adjacent to the master fault, sandstones are intercalated with, and diluted by, lacustrine or marine shales. in periods with nearshore depositional environments, fault trends acted either as a temporal or as a spatial controlling element on the sedimentary fill, but here the best reservoir quality is observed near the master fault, as seen, for example, in the gert area. the structural morphology is also relevant to prediction of the distribution of turbidite sandstones, which were probably deposited within lows on the hangingwall. sea-level changes in addition to the tectonic framework, sea-level variation is also believed to have controlled the location of sandstones. a sea-level curve for the middle jurassic to lower cretaceous succession in the danish central graben has been constructed by andsbjerg & dybkjær (2003, this volume). particularly low sea levels are interpreted to have occurred during the kimmeridgian; johannessen & andsbjerg (1993) indicated the presence of lowstand deposits associated with a sequence boundary in the ravn-1 well. this part of the jurassic succession, the heno formation, was deposited in the kimmeridgian during a rift relaxation phase and a likely explanation for the formation of this boundary is a regional (eustatic?) sea-level fall. sandstones may thus have been laid down in the basinal areas of the tail end graben and provide additional reservoir potential within the jurassic succession. the turbidites encountered in the jeppe-1 well were interpreted by johannessen et al. (1996) to represent lowstand deposits. the sea-level lowstand is suggested to have promoted the deposition of coarse-grained deposits in the basinal areas, which only received argillaceous sediments during sea-level highstand (plint 1988). the importance of sea-level changes is thought to be enhanced during the post-rift stage (nøttvedt et al. 1995) or during intra-rift periods of tectonic quiescence, as is the case for the heno formation and the turbidites encountered in the jeppe-1 well. in the middle jurassic – lowermost cretaceous succession of the danish central graben, the three tectonic phases documented here are recorded by discrete sedimentary units showing characteristic architectural features. in the sequence stratigraphic study of this succession by andsbjerg & dybkjær (2003, this volume), however, twelve sequences have been recognised, reflecting the interaction between the three major tectonic pulses described here and regional sea-level change. conclusions the danish central graben evolved through three distinct tectonic pulses. the first pulse, in callovian to early oxfordian times, was concentrated along preexisting n–s-trending faults. the second pulse occurred during the early volgian, along new nnw–sse-striking faults. most of the fault blocks were tilted towards the east. during this pulse, salt tectonics became important for the location of sedimentary depocentres which were displaced away from the master fault resulting in a more symmetrical configuration of subbasins. the third pulse was a continuation of the second pulse but accompanied by local thrusting. the structural framework controlled to some degree the distribution of reservoir sandstones. the most important potential reservoir facies related to periods with rotational tilt occur within the middle jurassic bryne and lulu formations, and the upper jurassic ‘fife sandstone formation’. reservoir sandstones related to tectonic relaxation are the heno formation and upper jurassic turbidites within the upper farsund formation. regional (eustatic?) sea-level variation may have been important in the latter case, especially with respect to deposition of sands in basinal areas. 262 263 acknowledgements the authors would like to thank claus andersen for preparing the geological maps and nina skårup for technical work on the geosections (figs 3–5). jan andsbjerg, karen dybkjær and ulrik gregersen are thanked for valuable comments on the manuscript. special thanks are due to john a. korstgård, arvid nøttvedt and jon r. ineson for reviewing the manuscript. references alexander, j. & leeder, m.r. 1987: active tectonic control on alluvial architecture. in: ethridge f.g., flores, r.m. & harvey, m.d. 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(ed.): petroleum geology of northwest europe: proceedings of the 4th conference, 1255–1267. london: geological society. surlyk, f. 1989: mid-mesozoic synrift turbidite systems: controls and prediction. in: collinson, j.d. (ed.): correlation in hydrocarbon exploration, 231–241. london: graham & trotman for the norwegian petroleum society (npf). underhill, j.r. & partington, m.a. 1993: jurassic thermal doming and deflation in the north sea: implications of the sequence stratigraphic evidence. in: parker, j.r. (ed.): petroleum geology of northwest europe: proceedings of the 4th conference, 337–346. london: geological society. vejbæk, o.v. 1986: seismic stratigraphy and tectonic evolution of the lower cretaceous of the danish central trough. danmarks geologiske undersøgelse serie a 11, 46 pp. vejbæk, o.v. 1992: geodynamic modelling of the danish central trough. in: larsen, r.m. et al. (eds): structural and tectonic modelling and its application to petroleum geology. norwegian petroleum society (npf) special publication 1, 1–17. ziegler, p.a. 1975: geologic evolution of the north sea and its tectonic framework. american association of petroleum geologists bulletin 59, 1073–1097. ziegler, p.a. 1982: geological atlas of western and central europe, 130 pp. the hague: elsevier for shell internationale petroleum maatschappij. manuscript received 30 june 1998; revision accepted 17 february 1999. geological survey of denmark and greenland bulletin 6, 1-7 1 geological survey of denmark and greenland bulletin 5 · 2004 the jurassic of north-east greenland edited by lars stemmerik and svend stouge geological survey of denmark and greenland ministry of the environment geus bulletin no 5.pmd 29-10-2004, 11:131 2 geological survey of denmark and greenland bulletin 5 keywords ammonites, boreal, dinoflagellate cysts, jurassic, north-east greenland, palaeogeography, rifting, siliciclastic sediments, stratigraphy cover eastwards-dipping middle–upper jurassic sandstones (yellow) and interbedded marine mudstones (dark) at the base of the coastal cliffs along the south-east coast of traill ø. the jurassic succession is described by vosgerau et al. (this volume). it is disconformably overlain by poorly exposed cretaceous siltstones with numerous volcanic intrusions that form ledges towards the top of the c. 1050 m high cliff. photo: lars stemmerik. frontispiece: facing page middle jurassic and lower cretaceous sandstones exposed on the western slopes of steensby bjerg, hold with hope, viewed towards the north-east with finsch øer in the centre and clavering ø in the far distance. on hold with hope, a more than 500 m thick sedimentary succession of triassic–cretaceous age is exposed in the north-facing coastal cliffs. the jurassic succession, which was not recognised until field work in 1996, is preserved in the downfaulted hangingwall blocks of a series of rotated half-grabens formed during the main east greenland rifting phase in the latest jurassic to earliest cretaceous. lower cretaceous sandstones, up to 170 m thick, unconformably overlie the rift succession. photo: michael larsen. chief editor of this series: peter r. dawes scientific editors: lars stemmerik and svend stouge, in conjunction with jon r. ineson copy editors: jon r. ineson and birgit eriksen editorial secretary: birgit eriksen critical readers: d.j. batten, g. bloos, walter k. christiansen, gregers dam, susanne feist-burkhardt, jon gjelberg, g.f.w. herngreen, jan jansonius, michael larsen, j.b. riding, d. strogen, finn surlyk, a. wierzbowski drawing work: jette halskov photographic work: jacob lautrup, benny m. schark lay-out and graphic production: knud gr@phic consult, odense, denmark printers: schultz grafisk, albertslund, denmark manuscripts submitted: 31 january 2000 – 20 march 2001 final versions approved: 22 january 2001 – 5 november 2002 printed: 1 november 2004 isbn 87-7871-135-5 geological survey of denmark and greenland bulletin the series geological survey of denmark and greenland bulletin replaces geology of denmark survey bulletin and geology of greenland survey bulletin. citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 5, 112 pp. available from geological survey of denmark and greenland øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk or geografforlaget aps rugårdsvej 55, dk-5000 odense c, denmark phone: +45 63 44 16 83, fax: +45 63 44 16 97, e-mail: go@geografforlaget.dk © danmarks og grønlands geologiske undersøgelse (geus), 2004 geus bulletin no 5.pmd 29-10-2004, 11:132 3 geus bulletin no 5.pmd 29-10-2004, 11:133 4 geus bulletin no 5.pmd 29-10-2004, 11:134 5 contents preface . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6 jurassic syn-rift sedimentation on a seawards-tilted fault block, traill ø, north-east greenland henrik vosgerau, peter alsen, ian d. carr, jens therkelsen, lars stemmerik and finn surlyk . . . . . . . . . . . . 9 the fluviatile bristol elv formation, a new middle jurassic lithostratigraphic unit from traill ø, north-east greenland jens therkelsen and finn surlyk . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19 maximum middle jurassic transgression in east greenland: evidence from new ammonite finds, bjørnedal, traill ø peter alsen and finn surlyk. appendix by john h. callomon . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 31 a new middle–upper jurassic succession on hold with hope, north-east greenland henrik vosgerau, michael larsen, stefan piasecki and jens therkelsen . . . . . . . . . . . . . . . . . . . . . . . . . . . 51 jurassic dinoflagellate cyst stratigraphy of hold with hope, north-east greenland stefan piasecki, michael larsen, jens therkelsen and henrik vosgerau . . . . . . . . . . . . . . . . . . . . . . . . . . . 73 jurassic dinoflagellate cysts from hochstetter forland, north-east greenland stefan piasecki and lars stemmerik . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 89 jurassic dinoflagellate cyst stratigraphy of store koldewey, north-east greenland stefan piasecki, john h. callomon and lars stemmerik . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 99 geus bulletin no 5.pmd 29-10-2004, 11:135 6 preface the jurassic sedimentary succession in east and northeast greenland reflects deposition during the early stages of rifting between greenland and norway. jurassic sediments are exposed over a distance of more than 600 km, from jameson land in the south to store koldewey in the north (fig. 1), and form one of the best-known exposed ancient rift successions. the sediments have been intensely studied over the last 25 years and a synthesis of the jurassic system in greenland was recently given in geological survey of denmark and greenland bulletin 1 (surlyk 2003). this collection of papers deals with stratigraphic and depositional aspects of the middle–upper jurassic sediments from isolated and less well-known localities outside the main outcrop areas, and thus adds to the tremendous amount of new data generated from the classical jurassic successions of jameson land and the wollaston forland area (ineson & surlyk 2003). most papers are based on fieldwork in 1996 and 1997, carried out within the framework of the project ‘resources of the sedimentary basins of north and east greenland’ supported by the danish research councils (see stemmerik et al. 1997). papers dealing with the jurassic at store koldewey and hochstetter forland are based on material collected during regional mapping in 1989 (stemmerik & piasecki 1990). during middle–late jurassic times, rifting took place along major n–s-trending synthetic faults that delimited wide westwards-tilted fault blocks (surlyk 1977, 2003). this resulted in the development of elongated marine embayments with major rivers entering from the north and dominantly axial sediment transport towards the south (surlyk 1978, 2003; engkilde & surlyk 2003). the jurassic syn-rift succession on south-eastern traill ø is an exception to this general pattern (vosgerau et al. 2004a, this volume). sedimentation took place on an eastwards-tilted fault block; the succession shows an eastwards proximal–distal decrease in sandstone–mudstone ratio, reflecting increasing water depths to the east. on the adjacent fault block to the west, a new lithostratigraphic unit, the bristol elv formation, has been erected to describe a succession of fluvio-lacustrine sediments at the base of the middle jurassic rift succession (therkelsen & surlyk 2004, this volume). the non-marine succession is overlain by shallow marine sandstones of the pelion formation (upper bajocian), succeeded in turn by 25–30 m of black silty mudstones of the fossilbjerget formation fig. 1. simplified geological map of east and north-east greenland showing the distribution of jurassic sediments. modified from surlyk (2003). c c' a 18°w 16°w 74°n 76°n 72°n 22°w 20°w saf pdmf df jurassic fault stauning alper fault post-devonian main fault dombjerg fault greenland milne land traill ø geographical society ø hold with hope hochstetter forland store koldewey wollaston forland pd m f sa f d f jameson land kong oscar fjord liverpool land 100 km 26°w 22°w24°w28°w 26°w 24°w geus bulletin no 5.pmd 29-10-2004, 11:136 7 (alsen & surlyk 2004, this volume). the presence of the fossilbjerget formation on southern traill ø indicates complete drowning of the sandy pelion system during maximum middle jurassic transgression (alsen & surlyk 2004, this volume). a new middle–upper jurassic succession was found in the hangingwalls of small fault blocks at hold with hope during fieldwork in 1996 (stemmerik et al. 1997). the up to 360 m thick succession and its stratigraphy are described in detail by vosgerau et al. (2004b, this volume) and piasecki et al. (2004a, this volume). the succession resembles that seen at wollaston forland and kuhn ø. the hold with hope area was flooded during late middle jurassic time; lower–middle callovian shallow marine sandstones of the pelion formation overlie lower triassic sediments (vosgerau et al. 2004b, this volume). the overlying sandstones of the payer dal formation are of middle–late oxfordian age. the uppermost part of the succession belongs to the bernbjerg formation. the youngest sediments are of late oxfordian – early kimmeridgian age based on dinoflagellate cysts (piasecki et al. 2004a, this volume). dinoflagellate cysts have also been used to date the scattered outcrops of middle–upper jurassic sediments at hochstetter forland and store koldewey further to the north (piasecki & stemmerik 2004, this volume; piasecki et al. 2004b, this volume). the dinoflagellate cyst assemblages of these northern outliers are readily correlated to assemblages described from the middle– upper jurassic further to the south in east greenland, and also show some resemblance to assemblages described from north greenland (piasecki et al. 2004b, this volume). lars stemmerik references alsen, p. & surlyk, f. 2004: maximum middle jurassic transgression in east greenland: evidence from new ammonite finds, bjørnedal, traill ø. in: stemmerik, l. & stouge, s. (eds): the jurassic of north-east greenland. geological survey of denmark and greenland bulletin 5, 31–49 (this volume). engkilde, m. & surlyk, f. 2003: shallow marine syn-rift sedimentation: middle jurassic pelion formation, jameson land, east greenland. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 813–863. ineson, j.r. & surlyk, f. (eds) 2003: the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 948 pp. piasecki, s. & stemmerik, l. 2004: jurassic dinoflagellate cysts from hochstetter forland, north-east greenland. in: stemmerik, l. & stouge, s. (eds): the jurassic of north-east greenland. geological survey of denmark and greenland bulletin 5, 89–97 (this volume). piasecki, s., larsen, m., therkelsen, j. & vosgerau, h. 2004a: jurassic dinoflagellate cyst stratigraphy of hold with hope, north-east greenland. in: stemmerik, l. & stouge, s. (eds): the jurassic of north-east greenland. geological survey of denmark and greenland bulletin 5, 73–88 (this volume). piasecki, s., callomon, j.h. & stemmerik, l. 2004b: jurassic dinoflagellate cyst stratigraphy of store koldewey, north-east greenland. in: stemmerik, l. & stouge, s. (eds): the jurassic of north-east greenland. geological survey of denmark and greenland bulletin 5, 99–112 (this volume). stemmerik, l. & piasecki, s. 1990: post-caledonian sediments in north-east greenland between 76° and 78°30′n. rapport grønlands geologiske undersøgelse 148, 123–126. stemmerik, l., clausen, o.r., korstgård, j., larsen, m., piasecki, s., seidler, l., surlyk, f. & therkelsen, j. 1997: petroleum geological investigations in east greenland: project ‘resources of the sedimentary basins of north and east greenland’. geology of greenland survey bulletin 176, 29–38. surlyk, f. 1977: stratigraphy, tectonics and palaeogeography of the jurassic sediments of the areas north of kong oscars fjord, east greenland. bulletin grønlands geologiske undersøgelse 123, 56 pp. surlyk, f. 1978: jurassic basin evolution of east greenland. nature 274(5667), 130–133. surlyk, f. 2003: the jurassic of east greenland: a sedimentary record of thermal subsidence, onset and culmination of rifting. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 659–722. therkelsen, j. & surlyk, f. 2004: the fluviatile bristol elv formation, a new middle jurassic lithostratigraphic unit from traill ø, north-east greenland. in: stemmerik, l. & stouge, s. (eds): the jurassic of north-east greenland. geological survey of denmark and greenland bulletin 5, 19–29 (this volume). vosgerau, h., alsen, p., carr, i.d., therkelsen, j., stemmerik, l. & surlyk, f. 2004a: jurassic syn-rift sedimentation on a seawards-tilted fault block, traill ø, north-east greenland. in: stemmerik, l. & stouge, s. (eds): the jurassic of north-east greenland. geological survey of denmark and greenland bulletin 5, 9–18 (this volume). vosgerau, h., larsen, m., piasecki, s. & therkelsen, j. 2004b: a new middle–upper jurassic succession on hold with hope, north-east greenland. in: stemmerik, l. & stouge, s. (eds): the jurassic of north-east greenland. geological survey of denmark and greenland bulletin 5, 51–71 (this volume). geus bulletin no 5.pmd 29-10-2004, 11:137 geological survey of denmark and greenland bulletin 1, 723-775 723 the aim of this study was to obtain a better understanding of the palynomorph flora, the age and the depositional environment of the neill klinter group in jameson land, east greenland (fig. 1). although the sedimentology, fossil faunas, ichnofaunas, lithostratigraphy and sequence stratigraphy of this succession have been studied in detail (rosenkrantz 1934; sykes 1974; dam 1990a, b, 1991; dam & surlyk 1995, 1998), there are few published papers on the palynology of the neill klinter group. the group has recently been divided into four formations and nine members (fig. 2) and a detailed sequence stratigraphic correlation between east greenland and norway has been established (dam & surlyk 1995, 1998). lithostratigraphic units of the group that are precisely dated by macrofossils are restricted to the rævekløft formation and the lepidopteriselv, palynostratigraphy and palaeoenvironments of the rævekløft, gule horn and ostreaelv formations (lower–middle jurassic), neill klinter group, jameson land, east greenland eva b. koppelhus and gregers dam the neill klinter group of jameson land, east greenland contains rich and diverse palynomorph assemblages. spores, pollen and freshwater algae dominate most of the samples, but dinoflagellate cysts and acritarchs also form important components. the ages suggested by the palynomorphs from the rævekløft, gule horn and ostreaelv formations span the period from the early pliensbachian to the early aalenian. the number of palynomorphs identified totals 136, including 83 miospore and 53 microplankton species; they are grouped into seven palynological assemblage zones. in general, there is good agreement between the palynological and sedimentological data, and the palynological data has refined the understanding of the depositional palaeoenvironments of the neill klinter group. in some cases, the boundaries of the palynological assemblage zones are congruent with major sequence stratigraphic surfaces and the palynological data thus support the sequence stratigraphic interpretation. in other cases, however, regional correlation indicates that the zone boundaries cross important sequence stratigraphic surfaces, such as sequence boundaries; such behaviour is thought to reflect the facies-dependent nature of certain of the palynological assemblage zones. the pattern of palynological events in east greenland has also been recognised on the mid-norwegian shelf. keywords: east greenland, jameson land basin, lower–middle jurassic, early pliensbachian – early aalenian, palynostratigraphy, sedimentology, sequence stratigraphic implications, regional correlation e.b.k.* & g.d.‡, geological survey of denmark and greenland, geocenter copenhagen, øster voldgade 10, dk1350 copenhagen k, denmark. present addresses: *royal tyrrell museum of palaeontology, box 7500, drumheller t0j 0y0, alberta, canada. email: evakoppelhus@hotmail.com ‡dong a/s, agern allé 24–26, dk-2970 hørsholm, denmark. geological survey of denmark and greenland bulletin 1, 723–775 (2003) © geus, 2003 724 nathorst fjeld and skævdal members of the ostreaelv formation. the neill klinter group was sampled intensively for palynological analysis at a number of different localities in jameson land (fig. 1). this paper only includes data from the lowermost three formations of the neill klinter group (rævekløft, gule horn and ostreaelv formations). the uppermost formation, the sortehat formation, is treated in an accompanying paper (koppelhus & hansen 2003, this volume). seven palynomorph assemblage zones have been established from the most complete section, at albuen (figs 3, 4). data from other localities have been correlated with this section. geological setting the upper palaeozoic – mesozoic jameson land basin is located in the present-day land areas of jameson land and scoresby land, at the southern end of the east greenland rift system (fig. 1; surlyk 1978). this system is part of a larger rift complex separating greenland from norway before the opening of the north atlantic ocean (ziegler 1988). the jameson land basin is bounded to the east and west by major n–s-trending faults, and to the north by a nw–se cross-fault in kong oscar fjord (surlyk 1977a, 1978, 1990a). the southern boundary is 25 km illoqqortoormiut 24°w 72°n 71°n n neill klinter group studied localities normal fault ? ? scoresby sund ? jameson land scoresby land kong o scar fjord liaselv rhætelv lepidopteriselv ranunkeldal enhjørningen dal horsedal moskusoksekløft astartekløft goniomyakløft albuen rævekløft harris fjeld nathorst fjeld elis bjerg dusén bjerg sortehat primulaelv qupaulakajik skævdal vardekløft tancrediakløft ? 22°w liverpool land greenland fig. 1. map of the jameson land region showing the outcrop of the neill klinter group, the location of sections forming the basis of this study and additional localities mentioned in the text. 725 unknown, but the basin probably extended south of scoresby sund, an area that is now covered by palaeogene plateau basalts. the basin was initiated in the devonian due to extensional collapse of the over-thickened crust of the caledonian mountain belt. the devonian phase was probably associated with strike-slip or oblique-slip deformation resulting in the development of nw–se-trending transverse faults in the north-eastern part of the basin. during late carboniferous – early permian times, the oblique-slip regime changed to a more orthogonal extensional regime, resulting in the development of basin margin half-grabens (surlyk et al. 1984, 1986; surlyk 1990a; larsen & marcussen 1992). the period of extensional tectonics was followed by a long period of subsidence lasting from the late permian to the cretaceous, interrupted by minor episodes of rifting and faulting (surlyk 1977a, b, 1990a; clemmensen 1980a; surlyk et al. 1981, 1986; larsen & marcussen 1992). triassic – early jurassic sedimentation in the jameson land basin was, in addition to tectonic subsidence, also influenced by climate, drainage patterns and eustasy. during the triassic – earliest jurassic, a major lacustrine complex was situated in the jameson land basin. the lacustrine deposits record a long-term change from a warm arid to a more temperate humid climate (clemmensen 1978a, b, 1979, 1980a, b; bromley & asgaard 1979; dam & surlyk 1992, 1993). this long-term climatic change was mainly governed by a gradual northwards drift of the laurasian continent and was accompanied by a long-term eustatic sea-level rise during the early jurassic. in the pliensbachian, the lacustrine complex was transformed into a shallow marine embayment marking the first fully-marine inundation of the basin since late permian – early triassic times (surlyk 1990b). the sandstones and mudstones of the neill klinter group were deposited in a wide, shallow tideand storm-influenced marine embayment, during a period of relative tectonic quiescence. the facies pattern was controlled mainly by relative sea-level fluctuations, sediment influx and basinal currents (sykes 1974; dam & surlyk 1995, 1998). sortehat ostreaelv gule horn elis bjerg skævdal harris fjeld nathorst fjeld lepidopteriselv trefjord bjerg memberformation albuen astartekløft horsedal rævekløft rhaetian hettangian sinemurian pliensbachian toarcian lo w er m id dl e ju ra ss ic tr ia s. u pp er aalenian bajocian bathonian callovian oxfordian kimmeridgian volgian groupchronostratigraphy formation sortehat ostreaelv gule horn rævekløft kap stewart ja m es on l an d supergroup scoresby sund hall bredning vardekløft neill klinter raukelv hareelv olympen fossilbjerget pelion r hæ te lvprimulaelv innakajik fig. 2. jurassic lithostratigraphy of jameson land showing the detailed subdivision of the neill klinter group formalised by dam & surlyk (1998). modified from surlyk (2003, this volume, fig. 5). stratigraphy the pliensbachian – lower aalenian succession now referred to the neill klinter group was initially described by rosenkrantz (1929), but was first formally established as a formation by surlyk et al. (1973). the neill klinter formation (sensu surlyk et al. 1973) has subsequently been raised to group status (dam & surlyk 1998) and the rævekløft, gule horn and ostreaelv members of surlyk et al. (1973) have been elevated to formation status. moreover, the former sortehat member of the vardekløft formation (surlyk et al. 1973) has been promoted to formation and transferred to the neill klinter group (dam & surlyk 1998). these authors divided the gule horn formation into two new members and the ostreaelv formation into seven new members (fig. 2). the neill klinter group is exposed in jameson land and scoresby land, and in a small fault-bounded outlier in the southern part of liverpool land (fig. 1); it is 300–450 m thick. the boundary between the kap stewart group and the succeeding neill klinter group is an erosional unconformity along the south-eastern basin margin, representing a major hiatus corresponding to the sinemurian stage (harris 1931; surlyk 1991; dam & surlyk 1995, 1998). the unconformity passes basinwards into a conformity and the contact between the lacustrine mudstones of the kap stewart group and the shallow marine sandstones of the neill klinter group is gradational (fig. 5; surlyk 1991; dam & surlyk 1995, 1998). the upper boundary of the neill klinter group is placed at a sharp unconformity between the mudstones of the sortehat formation and the sandstones of the vardekløft group (surlyk et al. 1973; surlyk 1990a; engkilde 1994; koppelhus & hansen 2003, this volume). the neill klinter group and most of its constituent formations and members show an overall sheet geometry, although the thicknesses of the units are greatest in the basin centre and thin towards the margins (dam & surlyk 1995, 1998). a rich marine fauna is present in the lower part of the neill klinter group (rosenkrantz 1934). it occurs in the rævekløft formation and is restricted to certain levels separated by largely unfossiliferous intervals. rosenkrantz (1934) identified a lower division, with a diverse fauna (150 species) dominated by bivalves, gastropods, cephalopods, echinoids and crinoids, and an upper division yielding a relatively sparse fauna (c. 20 molluscan species). ammonites of the genus uptonia occur in the lower division, suggesting that these beds belong to the early pliensbachian jamesoni zone (rosenkrantz 1934). in the upper division, rosenkrantz (1934) found two ammonites, beaniceras sp. and lytoceras fimbriatum. rosenkrantz (1934) referred this division to the ibex zone, although the bed also yielded an ammonite that appears to be aegoceras aff. capricornus of the maculatum group, indicative of the davoei zone (callomon 1961; surlyk et al. 1973). all the belemnites recovered from the rævekløft formation by rosenkrantz were apparently collected from the jamesoni zone interval. they indicate that the jamesoni zone as adopted by rosenkrantz (1934) includes the early pliensbachian jamesoni zone to at least the ibex zone and possibly the early davoie zone (doyle 1991). the marine macrofossils of the lepidopteriselv, nathorst fjeld, skævdal and trefjord bjerg members of the ostreaelv formation and the sortehat formation are bivalves, brachiopods, crinoids, belemnites, ammonites, and vertebrates (rosenkrantz 1934). ammonites collected on the top of elis bjerg from strata belonging to the lepidopteriselv member include dactylioceras semicelatum (simpson) sensu howarth 1992 (probably including d. groenlandicum rosenkrantz 1934) and hildaites sp. aff. murleyi (moxon). dactylioceras semicelatum belongs to the early toarcian tenuicostatum zone, semicelatum subzone (j.h. callomon, personal communication 1993). hildaites sp. is an early form, reminiscent of protogrammoceras. dactylioceras sp. has also been collected at nathorst fjeld in the nathorst fjeld member and in the lower part of the skævdal member, also suggesting an early toarcian tenuicostatum zone age (c. bjerrum and j.h. callomon, personal communications 1996). dactylioceras sp. and hildaites sp., suggestive of the early toarcian, have also been collected from the lepidopteriselv member in horsedal, although they were both loose specimens. phydoleoceras sp. has been collected on nathorst fjeld in the trefjord bjerg member, just beneath the boundary of the sortehat formation (c. bjerrum, personal communication 1996). the lepidopteriselv member has been correlated on sequence stratigraphic grounds with the nathorst fjeld member (dam & surlyk 1995, 1998). on nathorst fjeld, rosenkrantz (1934) collected a specimen of the belemnite parapassolotheuthis polita at an altitude of 494 m, and ‘parabrachybelus’ subaduncatus at 509 m. the lower level probably belongs to the nathorst fjeld member, and the upper level to the overlying skævdal member. the two species have restricted ranges and are not known to be widespread in europe. parapassolotheuthis polita has only been recorded from the early toarcian latest falciferum zone or earliest bifrons zone (commune subzone) in britain, while ‘para726 brachybelus’ subaduncatus, which so far has only been recorded from mainland europe, has a range probably restricted to the latest toarcian levesquei zone (doyle 1991). the ammonite dactylioceras semicelatum (simpson) has been collected at the base of the skævdal member at nathorst fjeld (c. bjerrum, personal communication 1996) and in the lepidopteriselv member on top of elis bjerg indicating an early toarcian tenuicostatum zone, semicelatum subzone age (j.h. callomon, personal communication 1993). based on these data, the nathorst fjeld and lepidopteriselv members and the lower part of the skævdal member include strata with an early toarcian tenuicostatum zone to latest falciferum zone or earliest bifrons zone age. belemnites suggest that the skævdal member may also include strata with a latest toarcian levesquei zone age (doyle 1991; dam & surlyk 1998), suggesting either that the skævdal member has a very long age range, that the belemnites cannot be used stratigraphically or that the d. semicelatum at the base of the skævdal member is reworked. dam & surlyk (1995, 1998) interpreted the neill klinter group within a sequence stratigraphic framework and attempted a sequence stratigraphic correlation with the coeval tilje, ror, ile and not formations on the midnorwegian shelf. this comparison demonstrated that the lower jurassic in both regions consists of six sequences and it appears feasible to directly correlate systems tracts on a scale of a few tens of metres between east greenland and the mid-norwegian shelf (see fig. 17). previous palynological work previous reports on the palynology of the neill klinter group have been published by lund & pedersen (1985) and underhill & partington (1994). the former authors studied the neill klinter group together with the overlying vardekløft group (sensu surlyk 2003, this volume, fig. 5) and the lower part of the hareelv formation. based on material collected from vardekløft, in the south-eastern part of the basin (fig. 1), lund & pedersen (1985) proposed four assemblage zones for the entire succession based on the miospore assemblages. dinoflagellate cysts were used to improve the age correlation of the spore-pollen assemblages. the three lowermost assemblage zones a, b and c of lund & pedersen (1985) cover the rævekløft, gule horn and ostreaelv formations of the neill klinter group. assemblage zone a is divided into subassemblages a1 and a2; the age of the zone was suggested to be late pliensbachian because of the presence of the spore kraeuselisporites reissingeri and the dinoflagellate cyst nannoceratopsis triceras. assemblage zone b is characterised by abundant spheripollenites subgranulatus and luehndea spinosa, and an early toarcian age was proposed. assemblage zone c is characterised by the incoming of the pollen callialasporites dampieri, the spores sestrosporites pseudoalveolatus and staplinisporites caminus and the dinoflagellate cyst parvocysta contracta (now susadinium scrofoides); these species were considered to indicate a late toarcian age for the lowermost assemblage zone c (subassemblage (c)). underhill & partington (1994) discussed the development of the lower jurassic in east greenland in connection with a sequence stratigraphic study of the north sea. they included 3 sections from jameson land, section 1 from liaselv, section 2 from vardekløft and section 3 from the harris fjeld/primulaelv area (fig. 1). sections 1 and 2 cover the uppermost few metres of the ostreaelv formation and all of the sortehat formation and vardekløft group. section 3 covers 170 m of the neill klinter group. underhill & partington (1994) analysed 48 samples and recognised 11 events. they suggested a late pliensbachian – earliest toarcian age for the gule horn formation and a toarcian age for the ostreaelv formation. materials and methods most of the samples used in this study are from the section at albuen, in the fifth ravine north of skævdal (figs 1, 3). intervals that proved inaccessible in this ravine were sampled at astartekløft (see figs 1, 9). samples were also obtained from rævekløft, tancrediakløft, qupaulakajik, albuen, goniomyakløft, astartekløft, moskusoksekløft, harris fjeld, primulaelv, lepidopteriselv, liaselv, horsedal and ranunkeldal (fig. 1). a number of samples collected by claus heinberg and tove birkelund in 1974 from lepidopteriselv were also included in the study. the samples were processed for their palynological content using the techniques adopted at the former geological survey of greenland, as described by nøhr-hansen (1993). over 210 samples were analysed for their palynological content by means of a transmission light microscope. two hundred specimens were counted in each sample and all species were registered in the range chart programme sis and on the video database at the geological survey of denmark and greenland, where the slides are stored. 727 728 285 280 275 405427 405426 265 270 405425 405424 341235 260 255 405423 341234 405422 m si pb sand 290 m 405430 bisaccate pollen dominate405429 341236 405428 st ac ke d tid al c ha nn el s st ac ke d tid al c ha nn el s el is b je rg m em be r a lb ue n m em be r a z 4 a ss em bl ag e z on e 3 l at e pl ie ns ba ch ia n sb4 w av e a nd st or m -d om . sh or ef ac e cerebropollenites thiergartii becomes rare m si pb sand 210 215 220 405403 405402 405401 225 230 405410 405405 405404 405406 341232 405408 235 240 405417 405416 405414 405413 405411 250 m 245 405421 405420 405419 405418 341233 r es tr ic te d s he lf su bt id al s an d sh ee t su bt id al s an d sh ee t st ac ke d tid al c ha nn el s w av e a nd st or m -d om . sh or ef ac e el is b je rg m em be r a ss em bl ag e z on e 1 ?e ar ly p lie ns ba ch ia n a ss em bl ag e z on e 2 a ss em bl ag e z on e 3 la te p lie ns ba ch ia n sb3 dinoflagellate cysts disappear kekryphalospora distincta nannoceratopsis senex and limbicysta bjaerkei abundant botryococcus and bisaccate pollen common lycopodiacidites rugulatus mancodinium semitabulatum parvocysta barbata albuen (a) 729 mudstone sandstone pebbly sandstone coal volcanic intrusive concretion siderised rip-up mudstone clasts/conglomerate sharp/erosive or irregular sharp/planar gradational parallel lamination lenticular bedding wavy bedding sedimentary features bed contacts flaser bedding planar cross-bedding trough cross-bedding cross-lamination incipient wave ripple lamination wave ripple cross-lamination hummocky and swaley cross-stratification coarse-grained ripples biota structureless structureless (with quartzite pebbles) rootlets plant fragments drifted plant stems/logs bivalves echinoderms belemnites weak moderate bioturbation intense arenicolites isp. diplocraterion parallelum gyrochorte comosa ophiomorpha nodosa phoebichnus trochoides monocraterion isp. helminthopsis isp. planolites beverleyensis taenidium serpentinum trace fossils cone-in-cone structures thalassinoides isp. unidentified sinuous horizontal burrow teichichnus isp. legend to sedimentary logs lithology conglomerate (qz, quartzite clasts) slumping cross-bedding with pebbles along foresets gastropods ammonites brachiopods crinoids curvolithos multiplex fig. 3a–c. sedimentological logs through the gule horn (a, b) and ostreaelv (b, c) formations, neill klinter group, at albuen (for location, see fig. 1); the gule horn formation comprises the elis bjerg and albuen members, the ostreaelv formation is composed of the astartekløft, nathorst fjeld, skævdal and trefjord bjerg members. sequence boundaries (sb3–7), palynomorph assemblage zones (az) and sample numbers are indicated. the legend accompanying this figure is also applicable to figs 5, 7, 9 and 15. 730 m si pb sand 315 320 sb5 325 330 335 340 345 350 355 360 m a st ar te kl øf t m em be r ea rl y t oa rc ia n n at ho rs t fj el d m em be r t id al c ha nn el t id al c ha nn el su bt id al s an d sh ee t st or m -d om in at ed sa nd y sh oa l m si pb sand 287 290 295 305 310 315 m 405434 few palynomorphs 405433 341239 poor assemblage – only bisaccate pollen 405432 405431 405430 405429 341236 405428 a lb ue n m em be r a ss em bl ag e z on e 4 la te p lie ns ba ch ia n st or m -d om in at ed o ffs ho re t ra ns iti on sb5 bisaccate pollen abundant botryococcus disappears intrusion a z 5 405466 albuen (b) 731 sk æ vd al m em be r a ss em bl ag e z on e 6 a z 7 la te t oa rc ia n ea rl y a al en ia n t re fjo rd b je rg m em be r so rt eh at fm drowning surface 440 m 430 435 420 425 410 415 400 405 390 395 m si pb sand su bt id al s an d sh ee t bi ot ur ba te d sh el f sb6 sb7 341248 397452 405449 341247 405450 405451 405452 405453 405454 405455 405456 nannoceratopsis gracilis abundant p. halosa common botryococcus reappear callialasporites dampieri pareodinia halosa ? m si pb sand 390 m 380 375 385 370 365 360 r es tr ic te d sh el f sh or ef ac e n at ho rs t fj el d m em be r ea rl y t oa rc ia n a ss em bl ag e z on e 5 a ss em bl ag e z on e 6 a st ar te kl øf t m em be r 405457 341243 405458 405459 341245 405460 405462 405464 405466 341241 perinopollenites elatoides acme drowning surface k. reissingerii s. subgranulatus rare s. subgranulatus abundant cerebropollenites macroverrucosus becomes abundant dinoflagellates reappear n. senex luehndea spinosa spheripollenites subgranulatus abundant qz albuen (c) palynological zones: definition nine assemblage zones have been recognised in the neill klinter group. the zones are based on the composition of the entire assemblage of spores, pollen, dinoflagellate cysts, acritarchs and freshwater algae. the assemblage zones are numbered 1–9 and named after the species which dominate the assemblages. assemblage zones 1–6 are defined and described in detail in this paper, together with a brief description of assemblage zone 7. assemblage zones 7–9 are formally defined in the accompanying paper on the sortehat formation (koppelhus & hansen 2003, this volume). the recorded taxa are listed in full in appendix 1 and the important species are illustrated in plates 1–7. the most complete section was sampled at albuen (figs 1, 3); approximately 70 samples were analysed for palynomorphs from this section (fig. 4, facing page 744). these samples yielded rich though not very wellpreserved palynomorph assemblages (plates 1–7), with the exception of samples taken close to the palaeogene sills and dykes that penetrate the sedimentary succession. these samples were either barren or the palynomorphs present were so dark as to be indeterminate. from the productive samples, approximately 150 species of spores, pollen, dinoflagellate cysts, acritarchs and freshwater algae were identified (appendix 1). the intensively sampled albuen section forms the reference section for assemblage zones 1–6 defined in this paper; assemblage zones 7–9 are defined from the borehole at sortehat (fig. 1; koppelhus & hansen 2003, this volume). in addition to the albuen reference section, important data for the definition of the assemblage zones were obtained from ranunkeldal, goniomyakløft and astartekløft (fig. 1). in addition to the formal assemblage zones defined here, a distinctive palynological assemblage was identified in samples from the horsedal member of the ostreaelv formation at horsedal. this assemblage is defined as the deltoidospora assemblage (see below under the horsedal locality). assemblage zone 1: cerebropollenites thiergartii – pinuspollenites minimus – botryococcus new assemblage zone occurrence. albuen 211–222.5 m goniomyakløft 226.5 m (single sample) primulaelv 310 m (single sample) rævekløft 118–122 m ranunkeldal 306–352.5 m this assemblage zone was recorded from the rævekløft formation at rævekløft and goniomyakløft and from the elis bjerg member (gule horn formation) at albuen, primulaelv and ranunkeldal. relative to the sequence stratigraphic scheme of dam & surlyk (1995, 1998), the assemblage occurs within sequences sq1 and sq2 (see fig. 17). reference section. albuen, 211 m (sample 405401) – 222.5 m (sample 405404; figs 1–4). additional section. ranunkeldal, 306 m (sample 341171) – 352.5 m (sample 341173; figs 5, 6). base. the base is not seen in the reference section at albuen (figs 3a, 4). at ranunkeldal, the base of the zone is placed at sample 341171, immediately above the base of the neill klinter group (figs 5, 6); samples from the underlying kap stewart group are considered to represent a different assemblage but further work is required to precisely define the base of assemblage zone 1. top. the top of the zone is defined by the last sample showing this assemblage (sample 405404 at 222.5 m at albuen) beneath the first appearance of nannoceratopsis senex and n. sp. (figs 3a, 4). characteristics. terrestrial palynomorphs (spores and pollen) dominate together with botryococcus sp. the most common spores are deltoidospora sp. and baculatisporites sp., and the pollen is dominated by cerebropollenites thiergartii and pinuspollenites minimus and other bisaccates. other characteristic species are cerebropollenites macroverrucosus, chasmatosporites hians and c. major. no dinoflagellate cysts have been determined with certainty, but a few acritarchs were identified. the freshwater alga botryococcus sp. occurs in abundance. suggested age. an early pliensbachian age is proposed based on the presence of abundant cerebropollenites thiergartii. palaeoenvironment. the assemblage records a strong terrestrial signal, most components being indicative of freshwater to brackish conditions. botryococcus, for 732 733 m pbsi sand m pbsi sand 285 290 295 341168 341167 341169 300 305 310 315 320 325 m 341171 341170 341172 330 335 340 345 350 355 360 m 341173 o pe n la cu st ri ne st or m -d om in at ed o ffs ho re t ra ns iti on k ap s te w ar t g ro up w av e a nd s to rm -d om in at ed s ho re fa ce el is b je rg m em be r (g ul e h or n fm ) el is b je rg m em be r a ss em bl ag e z on e 1 a ss em bl ag e no t de fin ed a ss em bl ag e z on e 1 t id al c ha nn el ranunkeldal sb1 qz fig. 5. sedimentological log through the uppermost part of the kap stewart group and the elis bjerg member (gule horn formation) of the neill klinter group in ranunkeldal (for location, see fig. 1). sequence boundary (sb1), palynomorph assemblage zone 1 and sample numbers are indicated. for legend, see fig. 3; arrows denote grainsize trends. 734 r an un ke ld al system lower jurassic stage lithostratigraphy kap stewart group not defined lower pliensbachian gule horn formation elis bjerg member sinemurian palynological assembl. zones (m) 35 0 34 0 30 0 29 0 31 0 32 0 sample height 35 2. 50 31 1. 00 30 6. 00 29 9. 00 29 3. 00 28 9. 00 28 5. 00 sample number 34 11 73 34 11 72 34 11 71 34 11 70 34 11 69 34 11 68 34 11 67 1trilete sp. 2ischyosporites variegatus 3striatella parva 4retitriletes clavatoides 5retitriletes sp. 6deltoidospora spp. 7baculatisporites sp. 8tripartina variabilis 9ischyosporites sp. 10conbaculatisporites mesozoicus 11retitriletes semimuris 12lycopodiacidites rugulatus 13rogalskaisporites cicatricosus 14manumia delcourtii 15calamospora tener 16retitriletes austroclavatoides 17cibotiumsporites jurienensis 18megaspore spp. 19densoisporites scanicus 20striatella jurassica 21densosporites variabilis 22kekryphalospora distincta 23sestrosporites pseudoalveolatus 24chasmatosporites major 25bisaccate spp. 26cerebropollenites thiergartii 27pinuspollenites minimus 28chasmatosporites apertus 29corollina torosus 30chasmatosporites hians 31perinopollenites elatoides 32quadraeculina anellaeformis 33cerebropollenites macroverrucosus 34ricciisporites tuberculatus 35vesicaspora fuscus 36vittatina sp. 37araucariacites australis 38ovalipollis ovalis 39dinocyst sp. 40nannoceratopsis gracilis 41veryhachium spp. 42veryhachium reductum 43acritarch spp. 44leiofusa jurassica 45micrhystridium exilium 46leiosphaeridia spp. 47micrhystridium lymensis 48botryococcus spp. 49tasmanites sp. 50foraminifera spp. r ? ? r ? r ? a lp ha be tic al s pe ci es li st 43 a cr ita rc h sp p. 37 a ra u ca ri a ci te s a u st ra lis 7 b a cu la ti sp or it es s p. 25 bi sa cc at e sp p. 48 b ot ry oc oc cu s sp p. 15 c a la m os p or a t en er 33 c er eb ro p ol le n it es m a cr ov er ru co su s 26 c er eb ro p ol le n it es t h ie rg a rt ii 28 c h a sm a to sp or it es a p er tu s 30 c h a sm a to sp or it es h ia n s 24 c h a sm a to sp or it es m a jo r 17 c ib ot iu m sp or it es j u ri en en si s 10 c on b a cu la ti sp or it es m es oz oi cu s 29 c or ol lin a t or os u s 6 d el to id os p or a s pp . 19 d en so is p or it es s ca n ic u s 21 d en so sp or it es v a ri a b ili s 39 d in oc ys t sp . 50 fo ra m in ife ra s pp . 9 is ch yo sp or it es s p. 2 is ch yo sp or it es v a ri eg a tu s 22 k ek ry p h a lo sp or a d is ti n ct a 44 le io fu sa j u ra ss ic a 46 le io sp h a er id ia s pp . 12 ly co p od ia ci d it es r u gu la tu s 14 m a n u m ia d el co u rt ii 18 m eg as po re s pp . 45 m ic rh ys tr id iu m e xi liu m 47 m ic rh ys tr id iu m l ym en si s 40 n a n n oc er a to p si s gr a ci lis 38 o va lip ol lis o va lis 31 pe ri n op ol le n it es e la to id es 27 pi n u sp ol le n it es m in im u s 32 q u a d ra ec u lin a a n el la ef or m is 16 r et it ri le te s a u st ro cl a va to id es 4 r et it ri le te s cl a va to id es 11 r et it ri le te s se m im u ri s 5 r et it ri le te s sp . 34 r ic ci is p or it es t u b er cu la tu s 13 r og a ls k a is p or it es c ic a tr ic os u s 23 s es tr os p or it es p se u d oa lv eo la tu s 20 s tr ia te lla j u ra ss ic a 3 s tr ia te lla p a rv a 49 ta sm a n it es s p. 1 tr ile te s p. 8 tr ip a rt in a v a ri a b ili s 42 ve ry h a ch iu m r ed u ct u m 41 ve ry h a ch iu m s pp . 35 ve si ca sp or a f u sc u s 36 v it ta ti n a s p. in te rv al n ot s am pl ed 1 fi g. 6 . p al yn o m o rp h d is tr ib u tio n c h ar t fo r th e u p p er m o st p ar t o f th e k ap s te w ar t g ro u p a n d t h e e lis b je rg m em b er ( g u le h o rn f o rm at io n ) o f th e n ei ll k lin te r g ro u p i n r an u n ke ld al (f o r lo ca tio n , se e fi g. 1 ). f o r le ge n d , se e fi g. 4 . example, is a planktonic green alga that occurs in colonies. it is known to adapt to different aquatic environments (fresh to brackish water), has been recorded from tropical to subarctic regions and has a stratigraphic range from the precambrian to the present day (guy-ohlson 1992). when large numbers of botryococcus are recorded, it indicates that the depositional environment was strongly influenced by fresh or brackish waters (guy-ohlson 1992). remarks. assemblage zone 1 is equivalent to assemblage zone a of lund & pedersen (1985), and probably to the lower part of section 3 of underhill & partington (1994). the assemblage is similar but not identical to that described from the chasmatosporites zone (miospore) and the mendicodinium reticulatum zone (dinoflagellate) from the successions of bornholm and the øresund area (koppelhus & nielsen 1994; koppelhus & batten 1996). assemblage zone 2: nannoceratopsis–botryococcus new assemblage zone occurrence. albuen 229–241 m goniomyakløft 229.8–235 m lepidopteriselv 638–654 m liaselv 293–325 m rævekløft 208–210 m this assemblage zone is characteristic of the elis bjerg member of the gule horn formation where it typically occurs within sequence sq3 of dam & surlyk (1995, 1998), for example at albuen (figs 3a, 4), lepidopteriselv, liaselv and rævekløft. at goniomyakløft, however, assemblage zone 2 was identified in the uppermost rævekløft formation and the lowermost elis bjerg member (figs 7, 8), within sequence sq2 of dam & surlyk (1995, 1998). reference section. albuen, 229 m (sample 405405) – 241 m (sample 405417; figs 3a, 4). additional section. goniomyakløft, 229.8 m (sample 405469) – 235 m (sample 405471; figs 7, 8). base. the base of the zone is defined by the sample showing the first appearance of nannoceratopsis senex and n. sp. top. the top of the zone is defined by the last sample showing this palynomorph assemblage, above which nannoceratopsis disappears together with most other dinoflagellate cysts. 735 240 245 m m si pb sand 225 230 235 405471 405470 405469 405468 405467 drowning surface goniomyakløft st or m -d om in at ed sa nd y sh oa l su bt id al s an d sh ee t r es tr ic te d sh el f u pp er s ho re fa ce r æ ve kl øf t fo rm at io n el is b je rg m em be r a ss em bl ag e z on e 2 g ul e h or n fo rm at io n sb3 az1 fig. 7. sedimentological log through the uppermost part of the rævekløft formation and the elis bjerg member (gule horn formation) at goniomyakløft (for location, see fig. 1). sequence boundary (sb3), palynomorph assemblage zones (az) and sample numbers are indicated. for legend, see fig. 3. note that the exact location of sample 405467 is uncertain due to imprecise field records; although here placed immediately beneath the nonexposed interval (226.5–227.8 m), it is possible that it derives from immediately above this interval. for legend, see fig. 3. 736 g on io m ya kl øf t lower pliensbachian rævekløft formation lower jurassic 23 5. 00 23 4. 50 23 2. 40 22 9. 80 22 6. 50 40 54 71 40 54 70 40 54 69 40 54 68 40 54 67 1kekryphalospora distincta 2retitriletes semimuris 3retitriletes austroclavatoides 4stereisporites stereoides 5lycopodiacidites rugulatus 6baculatisporites sp. 7deltoidospora spp. 8densoisporites velatus 9staplinisporites caminus 10retitriletes clavatoides 11striatella seebergensis 12leptolepidites sp. 13densoisporites scanicus 14retitriletes sp. 15foraminisporis jurassicus 16tigrisporites microrugulatus 17striatella scanica 18rogalskaisporites cicatricosus 19neoraistrickia gristhorpensis 20megaspore spp. 21striatella jurassica 22todisporites major 23cibotiumsporites jurienensis 24neoraistrickia sp. 25corollina torosus 26quadraeculina anellaeformis 27chasmatosporites apertus 28perinopollenites elatoides 29bisaccate spp. 30cerebropollenites thiergartii 31chasmatosporites hians 32cerebropollenites macroverrucosus 33chasmatosporites minor 34pinuspollenites minimus 35chasmatosporites major 36corollina sp. 37callialasporites sp. 38dapcodinium sp. 39mancodinium semitabulatum 40dinocyst spp. 41nannoceratopsis sp. 42crassosphaera spp. 43acritarch spp. 44tasmanites sp. 45botryococcus spp. 46foraminiferal linings 47foraminifera spp. ? ? ? ? ? ? ? ? a lp ha be tic al s pe ci es li st 43 a cr ita rc h sp p. 6 b a cu la ti sp or it es s p. 29 bi sa cc at e sp p. 45 b ot ry oc oc cu s sp p. 37 c a lli a la sp or it es s p. 32 c er eb ro p ol le n it es m a cr ov er ru co su s 30 c er eb ro p ol le n it es t h ie rg a rt ii 27 c h a sm a to sp or it es a p er tu s 31 c h a sm a to sp or it es h ia n s 35 c h a sm a to sp or it es m a jo r 33 c h a sm a to sp or it es m in or 23 c ib ot iu m sp or it es j u ri en en si s 36 c or ol lin a s p. 25 c or ol lin a t or os u s 42 c ra ss os p h a er a s p. 38 d a p co d in iu m s p. 7 d el to id os p or a s pp . 8 d en so is p or it es v el a tu s 13 d en so is p or it es s ca n ic u s 40 d in oc ys t sp p. 46 fo ra m in ife ra l l in in gs 47 fo ra m in ife ra s pp . 15 fo ra m in is p or is j u ra ss ic u s 1 k ek ry p h a lo sp or a d is ti n ct a 12 le p to le p id it es s p. 5 ly co p od ia ci d it es r u gu la tu s 39 m a n co d in iu m s em it a b u la tu m 20 m eg as po re s pp . 41 n a n n oc er a to p si s sp . 19 n eo ra is tr ic k ia gr is t h or p en si s 24 n eo ra is tr ic k ia s p. 28 pe ri n op ol le n it es e la to id es 34 pi n u sp ol le n it es m in im u s 26 q u a d ra ec u lin a a n el la ef or m is 3 r et it ri le te s a u st ro cl a va to id es 10 r et it ri le te s cl a va to id es 2 r et it ri le te s se m im u ri s 14 r et it ri le te s sp . 18 r og a ls k a is p or it es c ic a tr ic os u s 9 s ta p lin is p or it es c a m in u s 4 s te re is p or it es s te re oi d es 21 s tr ia te lla j u ra ss ic a 17 s tr ia te lla s ca n ic a 11 s tr ia te lla s ee b er ge n si s 44 ta sm a n it es s p. 16 t ig ri sp or it es m ic ro ru gu la tu s 22 to d is p or it es m a jo r system stage palynological assembl. zones lithostratigraphy elis bjerg mb upper pliensbach. (m) sample height sample number 22 5 23 5 21 5 20 5 12 fi g. 8 . p al yn o m o rp h d is tr ib u tio n c h ar t fo r th e u p p er m o st p ar t o f th e r æ ve kl ø ft f o rm at io n a n d t h e e lis b je rg m em b er ( g u le h o rn f o rm at io n ) at g o n io m ya kl ø ft ( fo r lo ca tio n , se e fi g. 1 ). fo r le ge n d , se e fi g. 4 . characteristics. terrestrial palynomorphs dominate the assemblage together with botryococcus sp., as in assemblage zone 1; the spore kekryphalospora distincta appears for the first time. the difference between this assemblage and that of assemblage zone 1 is the appearance of dinoflagellate cysts, including mancodinium semitabulatum, nannoceratopsis senex, n. gracilis, n. plegas, n. triangulata and parvocysta barbata and the presence of more acritarchs, including limbicysta bjaerkei. suggested age. a late pliensbachian age is suggested based on the first appearance of the spore kekryphalospora distincta in sample 405411 (236.50 m) in the albuen section (fig. 3a); this species is known to have a range from late pliensbachian to early bajocian (fenton & riding 1987). palaeoenvironment. although influenced by brackish to marine waters, there is still a strong terrestrial signal. remarks. the spore kekryphalospora distincta is also known from assemblages of pliensbachian age from the danish area (anholt borehole, øresund borehole 15 and the korsodde section on bornholm; seidenkrantz et al. 1993; koppelhus & nielsen 1994; koppelhus & batten 1996). the appearance of nannoceratopsis triangulata is particularly noteworthy as this species has previously only been recorded from nw germany where it occurs in the uppermost toarcian (prauss 1987). the appearance of limbicysta bjaerkei and parvocysta barbata is unexpected at this level, as they are known to have their first appearance in the bifrons zone (late early toarcian) in the north sea and svalbard and on the mid-norwegian shelf (bjærke 1980a; riding & thomas 1992; i. throndsen, personal communication 1996). however, these anomalous occurrences were also recognised at this level in the neill klinter group by underhill & partington (1994); l. bjaerkei was recorded in their section 3 between 360 m and 370 m. two explanations are possible: either p. barbata and l. bjaerkei have a longer range than recently reported or the sediments are younger than expected. in assemblage zone 2, acritarchs are more common than in the underlying zone and in the succeeding zones; this pattern is known from the lower jurassic in england and wales (wall 1965). assemblage zone 3: chasmatosporites – cerebropollenites thiergartii – botryococcus new assemblage zone occurrence. albuen 246–284.7 m astartekløft 266–328 m lepidopteriselv 674–700 m assemblage zone 3 is confined to the uppermost part of the elis bjerg member (gule horn formation) in the albuen and lepidopteriselv sections; relative to the sequence stratigraphic scheme of dam & surlyk (1995, 1998), the assemblage occurs within the upper levels of sequence sq3, below sequence boundary sb4 in these sections (fig. 3a). at astartekløft, however, the assemblage spans the boundary between the elis bjerg member and the succeeding albuen member, thus straddling the sequence boundary (sb4) between sequences sq3 and sq4 (fig. 10). it should be noted, however, that recognition of assemblage zone 3 is based on only two widely spaced samples at the astartekløft locality (fig. 10). reference section. albuen, 246 m (sample 405418) – 284.7 m (sample 405427; figs 3a, 4). additional section. astartekløft, 266 m (sample 346614) – 328 m (sample 346627; fig. 10) base. the base is placed at the first sample in which dinoflagellate cysts are absent or rare, succeeding samples of assemblage zone 2 characterised by a number of dinoflagellate species. top. the top of the assemblage is placed at the last sample showing the assemblage described below; above this level, the palynomorph assemblage is dominated by bisaccate pollen. characteristics. this zone is also dominated by terrestrial material. a number of spores have their first appearance, such as striatella jurassica, kraeuselisporites reissingeri, taurocusporites verrucatus and densoisporites velatus. the pollen species and botryococcus sp. are very consistent; only few dinoflagellate cysts and acritarchs were recorded. suggested age. a late pliensbachian age is proposed based on the absence of marker species indicative of 737 a younger age. the miospore assemblage is a continuation of assemblage zone 2. palaeoenvironment. the palynology displays an overwhelmingly terrestrial signal; there is very little evidence of marine influence. remarks. cerebropollenites thiergartii is common through much of the zone but becomes rare towards the top of the zone and in succeeding zones. assemblage zone 4: bisaccates new assemblage zone occurrence. albuen 287–297.8 m astartekløft 339–340 m this assemblage zone is restricted to the albuen member (gule horn formation) at both albuen and astartekløft (figs 3a, 3b, 4, 9, 10); it thus falls within sequence sq4 of dam & surlyk (1995, 1998). reference section. albuen, 287 m (sample 405428) – 297.8 m (sample 405434; figs 3a, 3b, 4). additional section. astartekløft, 339 m (sample 405472) – 340 m (sample 405473; figs 9, 10). base. the base is defined by the first sample dominated overwhelmingly by bisaccate pollen, to the exclusion of most other palynomorphs. top. the top of the zone is defined by the last sample composed predominantly of bisaccate pollen; the succeeding sample, defining the base of the overlying assemblage zone 5, is characterised by the first appearance of spheripollenites subgranulatus. characteristics. this assemblage zone is characterised by very poor preservation of the few palynomorphs present and by the absence of marine palynomorphs. bisaccate pollen are common in most of the samples. the spores, pollen and botryococcus sp. that were abundant in the assemblage zones 1–3 are absent in this zone. suggested age. a late pliensbachian age is assigned to this zone, as for assemblage zone 3 (see above). palaeoenvironment. taken at face value, the palynomorph data suggest that the sediments of the albuen member were deposited in a more distal position relative to the source than that suggested by the previous assemblage zones. amongst all palynomorphs, bisaccate pollen are known to be found farthest away from the source, because of their ability to be transported by air. however, bisaccate pollen are also known to have a thick wall and therefore may be preferentially preserved. thus, although lack of marine palynomorphs could be interpreted in terms of a non-marine environment, it could also have resulted from selective destruction of the more thin-walled marine palynomorphs during intrusion of palaeogene igneous sills and dykes in the albuen member. remarks. this interval was not recognised in previous studies by lund & pedersen (1985) and underhill & partington (1994). assemblage zone 5: spheripollenites subgranulatus – cerebropollenites macroverrucosus – luehndea spinosa new assemblage zone occurrence. albuen 359–375.5 m astartekløft 341–383 m moskusoksekløft 359 m (single sample) primulaelv 451 m (single sample) in the albuen reference section, assemblage zone 5 was only recorded from the lower nathorst fjeld member of the ostreaelv formation; it should be noted that the underlying astartekløft member (also ostreaelv formation) was not sampled in this section so the potential downwards range of the assemblage is poorly constrained (figs 3b, 3c, 4). at astartekløft itself, however, assemblage zone 5 extends from the lowermost astartekløft member up into the nathorst fjeld member (figs 9, 10). relative to the sequence stratigraphic scheme of dam & surlyk (1995, 1998), assemblage zone 5 occurs within sequence sq5, beginning immediately above the sequence boundary (sb5) in the astartekløft section (fig. 9) and extending up to some 10 m beneath the flooding surface at albuen (fig. 3c). reference section. albuen, 359 m (sample 405466) – 375.5 m (sample 405458; figs 3b, 3c, 4). 738 739 m si pb sand 335 345 355 365 m 375 drowning surface sb5 m si pb sand 385 m qz qz r es tr ic te d sh el f sh or ef ac e astartekløft su bt id al s an d sh ee t n at ho rs t fj el d m em be r a st ar te kl øf t m em be r a ss em bl ag e z on e 5 a z 4 a ss em bl ag e z on e 5 a lb ue n m em be r 341270 341269 405472 405473 405474 405475 405476 405477 405478 405483 405479 fig. 9. sedimentological log through the uppermost albuen member (gule horn formation) and the astartekløft and nathorst fjeld members (ostreaelv formation) at astartekløft (for location, see fig. 1). sequence boundary (sb5), palynomorph assemblage zones (az) and sample numbers are indicated. for legend, see fig. 3. 740 a st ar te kl ø ft lower jurassic toarcian ostreaelv formation astartekløft membernathorst fjeld member upper pliensbachian gule horn formation elis bjerg memberalbuen member 38 3. 00 37 2. 00 34 9. 00 34 5. 00 34 2. 00 34 1. 50 34 1. 00 34 0. 00 33 9. 00 32 8. 00 26 6. 00 34 12 70 34 12 69 40 54 79 40 54 77 40 54 76 40 54 75 40 54 74 40 54 73 40 54 72 34 66 27 34 66 14 1stereisporites stereoides 2densoisporites velatus 3retitriletes austroclavatoides 4conbaculatisporites mesozoicus 5calamospora tener 6lycopodiacidites rugulatus 7baculatisporites sp. 8deltoidospora spp. 9chomotriletes sp. 10rogalskaisporites cicatricosus 11kraeuselisporites reissingerii 12retitriletes sp. 13triletes sp. 14striatella jurassica 15ischyosporites variegatus 16foraminisporis jurassicus 17kekryphalospora distincta 18manumiadel courtii 19leptolepidites sp. 20striatella seebergensis 21cingulizonates inequalis 22cerebropollenites thiergartii 23chasmatosporites hians 24quadraeculina anellaeformis 25corollina torosus 26chasmatosporites major 27perinopollenites elatoides 28bisaccate spp. 29pinuspollenites minimus 30chasmatosporites apertus 31vesicaspora fuscus 32spheripollenites subgranulatus 33cerebropollenites macroverrucosus 34chasmatosporites sp. 35corollina meyeriana 36striate spp. 37taeniasporites rhaeticus 38nannoceratopsis senex 39nannoceratopsis plegas 40mancodinium semitabulatum 41nannoceratopsis gracilis 42nannoceratopsis triangulata 43mendicodinium reticulatum 44kallosphaeridium sp. 45nannoceratopsis triceras 46leiosphaeridia spp. 47micrhystridium spp. 48botryococcus spp. ? r r r a lp ha be tic al s pe ci es li st 7 b a cu la ti sp or it es s p. 28 bi sa cc at e sp p. 48 b ot ry oc oc cu s sp p. 5 c a la m os p or a t en er 33 c er eb ro p ol le n it es m a cr ov er ru co su s 22 c er eb ro p ol le n it es t h ie rg a rt ii 30 c h a sm a to sp or it es a p er tu s 23 c h a sm a to sp or it es h ia n s 26 c h a sm a to sp or it es m a jo r 34 c h a sm a to sp or it es s p. 9 c h om ot ri le te s sp . 21 c in gu liz on a te s in eq u a lis 4 c on b a cu la ti sp or it es m es oz oi cu s 35 c or ol lin a m ey er ia n a 25 c or ol lin a t or os u s 8 d el to id os p or a s pp . 2 d en so is p or it es v el a tu s 16 fo ra m in is p or is j u ra ss ic u s 15 is ch yo sp or it es v a ri eg a tu s 44 k a llo sp h a er id iu m s p. 17 k ek ry p h a lo sp or a d is ti n ct a 11 k ra eu se lis p or it es r ei ss in ge ri i 46 le io sp h a er id ia s pp . 19 le p to le p id it es s p. 6 ly co p od ia ci d it es r u gu la tu s 40 m a n co d in iu m s em it a b u la tu m 18 m a n u m ia d el co u rt ii 43 m en d ic od in iu m r et ic u la tu m 47 m ic rh ys tr id iu m s pp . 39 n a n n oc er a to p si s p le ga s 41 n a n n oc er a to p si s gr a ci lis 38 n a n n oc er a to p si s se n ex 42 n a n n oc er a to p si s tr ia n gu la ta 45 n a n n oc er a to p si s tr ic er a s 27 pe ri n op ol le n it es e la to id es 29 pi n u sp ol le n it es m in im u s 24 q u a d ra ec u lin a a n el la ef or m is 3 r et it ri le te s a u st ro cl a va to id es 12 r et it ri le te s sp . 10 r og a ls k a is p or it es c ic a tr ic os u s 32 s p h er ip ol le n it es s u b gr a n u la tu s 1 s te re is p or it es s te re oi d es 36 st ri at e sp p. 14 s tr ia te lla j u ra ss ic a 20 s tr ia te lla s ee b er ge n si s 37 ta en ia sp or it es r h a et ic u s 13 tr ile te s sp . 31 ve si ca sp or a f u sc u s in te rv al n ot s am pl ed in te rv al n ot s am pl ed system stage palynological assembl. zones stratigraphy (m) sample height sample number 26 5 27 0 32 5 33 0 33 5 34 0 34 5 35 0 37 5 38 0 37 0 45 3 u nc er ta in de te rm in at io n ve ry r ar e r ar e fe w c om m on a bu nd an t ? r fi g. 1 0. p al yn o m o rp h d is tr ib u tio n c h ar t fo r th e e lis b je rg a n d a lb u en m em b er s (g u le h o rn f o rm at io n ) an d t h e a st ar te kl ø ft a n d n at h o rs t fj el d m em b er s (o st re ae lv f o rm at io n ) at a st ar te kl ø ft ( fo r lo ca tio n , se e fi g. 1 ). additional section. astartekløft, 341 m (sample 405474) – 383 m (sample 341270; figs 9, 10). base. the base of the assemblage is defined by the first appearance, in abundance, of spheripollenites subgranulatus; at albuen, this coincides with the first appearance of luehndea spinosa, although this species was not recorded in the astartekløft section. top. this is defined by the uppermost sample showing the palynomorph assemblage characteristic of the zone (i.e. sample 405458 at albuen). the succeeding sample shows an acme of perinopollenites elatoides, defining the base of assemblage zone 6. the top of the assemblage is not seen in the astartekløft section (fig. 9). characteristics. in assemblage zone 5, the pollen spheripollenites subgranulatus appears in abundance for the first time; this species was only recorded in one sample below this level, in sample 346627 at 328 m in the astartekløft section (figs 9, 10). the assemblage is also marked by the reappearance of abundant nannoceratopsis senex and n. gracilis. the pollen cerebropollenites macroverrucosus and to a lesser degree corollina torosus are more common than in the underlying assemblage. spherical bodies of uncertain affinity are also abundant (on the distribution chart they are registered under micromiscellanea). suggested age. an early–late toarcian age is proposed based on the presence and acme of spheripollenites subgranulatus which is known to be abundant in toarcian strata in the danish basin and the north sea (dybkjær 1991; batten et al. 1994; koppelhus & nielsen 1994; koppelhus & batten 1996). palaeoenvironment. this assemblage contains elements indicative of both brackish and marine conditions. remarks. in germany, the north sea area and the danish basin, the dinoflagellate cyst luehndea spinosa is known to appear within the margaritatus, spinatus and tenuicostatum zones which span the late pliensbachian to early toarcian (morgenroth 1970; riding & thomas 1992; poulsen 1996). the abundant spherical bodies of unknown affinity are also known from svalbard, the north sea and the baltic sea in upper pliensbachian and toarcian strata (bjærke 1980b; dybkjær 1991; koppelhus & nielsen 1994); on the mid-norway shelf, these forms are known to occur in lower toarcian strata (i. throndsen, personal communication 1996). this zone correlates with assemblage zone b of lund & pedersen (1985). it has not been possible to recognise this assemblage in the data presented by underhill & partington (1994). comparison with the microcysta erugata taxa range-zone of smelror & below (1992) has been attempted, but the two zones have very few species in common. assemblage zone 6: perinopollenites elatoides new assemblage zone occurrence. albuen 377.5–434 m enhjørningen dal 414–424.67 m sortehat (core) 12.65–26.28 m in the albuen section, assemblage zone 6 extends from the upper nathorst fjeld member through the skævdal member and much of the trefjord bjerg member (all ostreaelv formation; figs 3c, 4). relative to the sequence stratigraphic scheme of dam & surlyk (1995, 1998), the assemblage occurs within sequences sq5 and sq6, spanning the sequence boundary sb6 (figs 3c, 4). data for the enhjørningen dal and sortehat sections is presented in the companion paper by koppelhus & hansen (2003, this volume). in the sortehat core, assemblage zone 6 is confined to the upper ostreaelv formation (trefjord bjerg member); the top of the assemblage lies immediately beneath the boundary between the ostreaelv and sortehat formations. at enhjørningen dal, however, assemblage zone 6 spans the boundary between these two formations, extending some 6 m up into the sortehat formation (koppelhus & hansen 2003, this volume). reference section. albuen, 377.5 m (sample 341243) – 434 m (sample 405449; figs 3c, 4). base. the base of the zone is placed at the sample in which perinopollenites elatoides reaches its acme; it is accompanied by abundant chasmatosporites hians and c. major. top. the upper boundary of assemblage zone 6 is placed at the last sample showing the palynomorph assemblage described here; above this level, botryococcus sp. becomes very dominant, defining the base of assemblage zone 7. 741 characteristics. the zone is characterised particularly by the acme of perinopollenites elatoides. the spores staplinisporites caminus and sestrosporites pseudoalveolatus have their first appearance within the zone together with the dinoflagellate cysts phallocysta eumekes, wallodinium laganum, scriniocassis sp. and dissiliodinium sp. the dinoflagellate cysts nannoceratopsis gracilis and n. senex continue to be common in most of the samples. botryococcus sp. re-appears at 392 m (sample 405456; fig. 3c) and continues to be common to the top of this zone. suggested age. a late toarcian – early aalenian age is suggested based on the first appearances of the spores staplinisporites caminus and sestrosporites pseudoalveolatus together with the dinoflagellate cysts phallocysta eumekes, wallodinium laganum, scriniocassis sp. and dissiliodinium sp. the occurrence of abundant perinopollenites elatoides is a feature of aalenian sediments in the danish basin and on bornholm (dybkjær 1991; koppelhus & nielsen 1994). palaeoenvironment. the palynomorphs indicate that the palaeoenvironment was influenced by both fresh, brackish and marine waters, suggesting a nearshore environment. remarks. the two species chasmatosporites hians and c. major are very abundant in the lowermost sample in this interval, higher up they become rare. in offshore mid-norway, c. hians has a maximum appearance after the spheripollenites acme (i. throndsen, personal communication 1996). the pollen callialasporites dampieri is known to appear in the uppermost toarcian and lowermost aalenian in north-west scotland (riding et al. 1991). this zone correlates with assemblage zone c in lund & pedersen (1985) and the assemblage found in sample b2/57 at approximately 460 m in primulaelv by underhill & partington (1994). the dinoflagellate cyst wallodinium laganum appears for the first time in this zone and is only known from the late toarcian levesquei zone in northern germany and england (feist-burkhardt & monteil 1994). the ammonite dactylioceras sp. occurs at the base of the skævdal member at nathorst fjeld, suggesting an early toarcian tenuicostatum zone age for the lower part of the assemblage zone. assemblage zone 7: botryococcus this assemblage is formally defined in the companion paper by koppelhus & hansen (2003, this volume); a summary is given here. occurrence. albuen 438.5–443.5 m enhjørningen dal 424.86–445 m pelion 550–567 m sortehat (core) 27.82–36.36 m at albuen, this assemblage is represented in the uppermost few metres of the trefjord member (ostreaelv formation) and extends up into the sortehat formation (figs 3c, 4); additional data for the uppermost trefjord member and the overlying sortehat formation in the albuen section are given in koppelhus & hansen (2003, this volume). in the cored section from sortehat, the assemblage occurs in the lower levels of the sortehat formation, the base of the zone being immediately above the base of the formation (koppelhus & hansen 2003, this volume). at enhjørningen dal, assemblage zone 7 is also restricted to the lower sortehat formation although here the base is some 6 m above the lower boundary of the sortehat formation. detailed discussion and the full dataset are given in koppelhus & hansen (2003, this volume). assemblage zone 7 occurs within the lower levels of sequence sq7 of dam & surlyk (1995, 1998). reference section. sortehat (core), 27.82 m (sample 303143-73) – 36.36 m (sample 303143-62). additional sections. see koppelhus & hansen (2003, this volume). base. the base of the assemblage is placed at the first sample in which botryococcus sp. overwhelmingly dominates the assemblage. in the albuen section (figs 3c, 4), this event coincides with the first co-occurrence of callialasporites dampieri (pollen) and mendicodinium groenlandicum (dinoflagellate cyst) although in other sections (e.g. enhjørningen dal, sortehat; koppelhus & hansen 2003, this volume) these species first occur together some metres below the botryococcus sp. influx. top. the upper boundary is defined by the uppermost sample showing the botryococcus-dominated assemblage. above this level, botryococcus sp. disappear and nannoceratopsis gracilis and n. senex become abundant once more. 742 characteristics. the assemblage is characterised by the overwhelming dominance of botryococcus sp. and the scarcity of dinoflagellates. suggested age. an aalenian age is proposed based on the abundance of callialasporites dampieri. pollen from the genus callialasporites are known to appear first in sediments of late toarcian and aalenian age in sweden and the danish area (guy-ohlson 1988; koppelhus & nielsen 1994). palaeoenvironment. the fresh and brackish water alga botryococcus is known from recent environments to produce blooms at certain times of the year. the colonies float at the water surface under calm conditions and subsequently sink when the water is disturbed. when they die, they float within the surface waters and can be transported by wind far from the area where they were produced. palaeoenvironmental interpretation based solely on the presence of botryococcus is therefore dangerous; the degree to which the botryococcus in this succession is allochthonous is unknown. further discussion of the environmental implications of this assemblage is given by koppelhus & hansen (2003, this volume). additional palynological results as noted earlier, definition of the six assemblage zones described here is based primarily on the section at albuen, the palynostratigraphy of which is thus presented in detail above. in addition to this reference section, however, a series of other sections were included in the study, some of which yield important supplementary data for the definition of the assemblage zones (see above). the palynostratigraphic results from these additional localities, spread widely in the jameson land basin (fig. 1), are described below, broadly from south to north. rævekløft at rævekløft, nine samples were collected, six from the rævekløft formation (405435–405440) and three from the elis bjerg member (gule horn formation, 405441–405443; fig. 11). the samples collected from the rævekløft formation are separated by a gap of c. 100 m from those collected in the elis bjerg member and the boundary between the two units was not exposed. all the samples yielded abundant poorly preserved palynomorphs. terrestrial material dominates together with the freshwater alga botryococcus sp. bisaccate pollen are the most abundant palynomorphs in all the samples. a few acritarchs and questionable dinoflagellate cysts were found. based on the presence of the spores deltoidospora and baculatisporites, the pollen cerebropollenites thiergartii and pinuspollenites minimus and the lack of dinoflagellate cysts, the samples 405435–405440 (?upper rævekløft formation) are assigned to assemblage zone 1. in the samples 405441–405443 (elis bjerg member), the dinoflagellate cysts nannoceratopsis senex and mancodinium semitabulatum appear for the first time, indicating that the assemblage belongs to assemblage zone 2. suggested age. an early pliensbachian age is suggested for assemblage zone 1, based on the presence of cerebropollenites thiergartii together with pinuspollenites minimus; a late pliensbachian age is proposed for assemblage zone 2 based on the occurrences of nannoceratopsis senex and mancodinium semitabulatum. tancrediakløft only one sample (341229; rævekløft formation) was studied from this locality. the sample contained only black material which was not identifiable. qupaulakajik only one sample (341254; albuen member, gule horn formation) was studied from this locality. the palynomorphs were black and indeterminate. goniomyakløft five samples were studied (figs 7, 8). the sample 405467, from the uppermost part of rævekløft formation, yielded an assemblage rich in poorly preserved palynomorphs, dominated by terrestrial material. however, several specimens of the dinoflagellate cyst genus dapcodinium were found together with a single specimen of tasmanites. the former are similar to dapcodinium priscum, but not identical to specimens of this species described from northwest europe. the assemblage is suggested to belong to assemblage zone 1. 743 744 lower jurassic upper pliensbachian gule horn formation 12 elis bjerg member lower pliensbachian rævekløft fm 21 0. 00 20 9. 00 20 8. 00 12 2. 00 12 1. 00 12 0. 00 11 9. 00 11 8. 50 11 8. 00 40 54 43 40 54 42 40 54 41 40 54 40 40 54 39 40 54 38 40 54 37 40 54 36 40 54 35 1rogalskaisporites cicatricosus 2conbaculatisporites mesozoicus 3lycopodiacidites rugulatus 4deltoidospora spp. 5retitriletes clavatoides 6retitriletes semimuris 7retitriletes sp. 8baculatisporites sp. 9tigrisporites microrugulatus 10foraminisporis jurassicus 11kraeuselisporites reissingeri 12tripartina variabilis 13stereisporites stereoides 14densoisporites scanicus 15retitriletes austroclavatoides 16todisporites minor 17kekryphalospora distincta 18cibotiumsporites jurienensis 19striatella parva 20todisporites major 21striatella jurassica 22chasmatosporites hians 23cerebropollenites thiergartii 24vesicaspora fuscus 25quadraeculina anellaeformis 26perinopollenites elatoides 27pinuspollenites minimus 28bisaccate spp. 29cerebropollenites macroverrucosus 30chasmatosporites major 31chasmatosporites apertus 32corollina torosus 33monosulcites punctatus 34taeniasporites rhaeticus 35nannoceratopsis senex 36acritarch spp. 37veryhachium sp. 38leiofusa jurassica 39leiosphaeridia spp. 40botryococcus spp. 41tasmanites sp. r a lp ha be tic al s pe ci es li st 36 a cr ita rc h sp p. 8 b a cu la ti sp or it es s p. 28 bi sa cc at e sp p. 40 b ot ry oc oc cu s sp p. 29 c er eb ro p ol le n it es m a cr ov er ru co su s 23 c er eb ro p ol le n it es t h ie rg a rt ii 31 c h a sm a to sp or it es a p er tu s 22 c h a sm a to sp or it es h ia n s 30 c h a sm a to sp or it es m a jo r 18 c ib ot iu m sp or it es j u ri en en si s 2 c on b a cu la ti sp or it es m es oz oi cu s 32 c or ol lin a t or os u s 4 d el to id os p or a s pp . 14 d en so is p or it es s ca n ic u s 10 fo ra m in is p or is j u ra ss ic u s 17 k ek ry p h a lo sp or a d is ti n ct a 11 k ra eu se lis p or it es r ei ss in ge ri 38 le io fu sa j u ra ss ic a 39 le io sp h a er id ia s pp . 3 ly co p od ia ci d it es r u gu la tu s 33 m on os u lc it es p u n ct a tu s 35 n a n n oc er a to p si s se n ex 26 pe ri n op ol le n it es e la to id es 27 pi n u sp ol le n it es m in im u s 25 q u a d ra ec u lin a a n el la ef or m is 15 r et it ri le te s a u st ro cl a va to id es 5 r et it ri le te s cl a va to id es 6 r et it ri le te s se m im u ri s 7 r et it ri le te s sp . 1 r og a ls k a is p or it es c ic a tr ic os u s 13 s te re is p or it es s te re oi d es 21 s tr ia te lla j u ra ss ic a 19 s tr ia te lla p a rv a 34 ta en ia sp or it es r h a et ic u s 41 ta sm a n it es s p. 9 t ig ri sp or it es m ic ro ru gu la tu s 20 to d is p or it es m a jo r 16 to d is p or it es m in or 12 tr ip a rt in a v a ri a b ili s 37 ve ry h a ch iu m s p. 24 ve si ca sp or a f u sc u s r æ ve kl øf t 20 0 20 5 21 0 12 5 12 0 11 5 system stage palynological assembl. zones lithostratigraphy (m) sample height sample number in te rv al n ot s am pl ed fi g. 1 1. p al yn o m o rp h d is tr ib u tio n c h ar t fo r th e r æ ve kl ø ft f o rm at io n a n d g u le h o rn f o rm at io n ( e lis b je rg m em b er ) at r æ ve kl ø ft ( fo r lo ca tio n , se e fi g. 1 ). f o r le ge n d , se e fi g. 4 . sample 405468 was collected just beneath the boundary between the rævekløft formation and the succeeding elis bjerg member (gule horn formation). this sample yielded a similar palynomorph assemblage to that described above but included the first appearance of the dinoflagellate cyst mancodinium semitabulatum. it is suggestive of the somewhat younger assemblage zone 2. the samples 405469–405471 are from the elis bjerg member. they lack recognisable dinoflagellate cysts although acritarchs and some questionable dinoflagellate cysts were found together with crassosphaera sp., foraminiferal inner linings and abundant botryococcus sp. this assemblage is also suggested to belong to assemblage zone 2. suggested age. an ?early–late pliensbachian age is proposed, based on the appearance of mancodinium semitabulatum. astartekløft fourteen samples from the northern and southern side of astartekløft were investigated palynologically (figs 9, 10; note that the barren samples are not shown on fig. 10). the lowermost samples, 346614 at 266 m and 346627 at 328 m, are from the lower elis bjerg member (gule horn formation) and the lower albuen member (gule horn formation), respectively (fig. 10). these samples are dominated by botryococcus sp. and bisaccate pollen, together with common cerebropollenites thiergartii and chasmatosporites hians, and are referred to assemblage zone 3. the two samples 405472 and 405473 from the upper albuen member (figs 9, 10) contain abundant bisaccate pollen; this and the lack of other palynomorphs indicate assemblage zone 4. samples 405474 and 405475 are from the lowermost beds of the astartekløft member; they record the first appearance of spheripollenites subgranulatus, together with abundant cerebropollenites macroverrucosus and the re-appearance of the dinoflagellate cyst nannoceratopsis senex. sample 405477, also from the astartekløft member, lacks spheripollenites subgranulatus but nannoceratopsis senex is common. in sample 405483 from the astartekløft member, the organic material is black and indeterminate and thus this sample does not appear on figure 10. in sample 341269, from the nathorst fjeld member, spheripollenites subgranulatus is abundant, and in the uppermost sample (341270), also from the nathorst fjeld member, spheripollenites subgranulatus is absent but nannoceratopsis senex is abundant together with cerebropollenites macroverrucosus. it is suggested that the assemblages recorded between sample 405474, at the base of the astartekløft member, and the uppermost sample 341270, in the nathorst fjeld member, are compatible with assemblage zone 5 as defined from the albuen section. suggested age. the lowermost part of the succession is referred to the upper pliensbachian based on the abundance of cerebropollenites thiergartii and chasmatosporites hians. an early toarcian age for the uppermost part is based on the first appearance and abundance of both spheripollenites subgranulatus and cerebropollenites macroverrucosus. moskusoksekløft only one sample (341260; nathorst fjeld member, ostreaelv formation) was investigated from this locality; it yielded the dinoflagellate cysts nannoceratopsis senex and n. triangulata and the pollen spheripollenites subgranulatus and cerebropollenites macroverrucosus (fig. 12a). this assemblage is referred to assemblage zone 5. suggested age. an early toarcian age is proposed based on the presence of spheripollenites subgranulatus and cerebropollenites macroverrucosus together with nannoceratopsis senex and n. triangulata. harris fjeld one sample (346741) was studied from this locality, from the lower part of elis bjerg member (fig. 12b). the palynological assemblage is dominated by poorly preserved bisaccate pollen and the assemblage was deemed too poor to determine to which assemblage zone it belongs. primulaelv three samples (346746, 346745, 346753) were analysed from this locality (fig. 12c), the first two from the elis bjerg member and the third from the lowermost mudstones of the skævdal member. the lowermost sample is tentatively referred to assemblage zone 1, based on a very poor assemblage of bisaccates, pinuspollenites 745 746 sy st em lo w er ju ra ss ic st ag e to ar ci an pa ly no lo gi ca l a ss em bl . z on es li th os tr at ig ra ph y n at ho rs t fj el d m b o st re ae lv f or m at io n sa m pl e he ig ht 359.00 sa m pl e nu m be r 341260 1 d el to id os p or a sp p. 2 st ri at el la ju ra ss ic a 3 is ch yo sp or ite s va ri eg at us 4 b ac ul at is p or ite s sp . 5 sp he ri p ol le ni te s su bg ra nu la tu s 6 c er eb ro p ol le ni te s m ac ro ve rr uc os us 7 c or ol lin a to ro su s 8 bi sa cc at e sp p. 9 c or ol lin a m ey er ia na 10 c ha sm at os p or ite s m aj or 11 c er eb ro p ol le ni te s th ie rg ar tii 12 n an no ce ra to p si s se ne x 13 n an no ce ra to p si s tr ia ng ul at a 14 c ym at io sp ha er a sp . 15 fo ra m in ife ra s pp . 16 b ot ry oc oc cu s sp p. alphabetical species list 4 baculatisporites sp. 8 bisaccate spp. 16 botryococcus spp. 6 cerebropollenites macroverrucosus 11 cerebropollenites thiergartii 10 chasmatosporites major 9 corollina meyeriana 7 corollina torosus 14 cymatiosphaera sp. 1 deltoidospora spp. 15 foraminifera spp. 3 ischyosporites variegatus 12 nannoceratopsis senex 13 nannoceratopsis triangulata 5 spheripollenites subgranulatus 2 striatella jurassica a: moskusoksekløft 5 b: harris fjeld sy st em ? st ag e li th os tr at ig ra ph y ? pa ly no lo gi ca l a ss em bl . z on es ? el is b je rg m b g ul e h or n fm sa m pl e he ig ht 269.00 sa m pl e nu m be r 346741 1 b a cu la ti sp or it es s p. 2 ly co p od ia ci d it es r u gu la tu s 3 d el to id os p or a s pp 4 r og a ls k a is p or it es c ic a tr ic os u s 5 bi sa cc at e po lle n 6 c h a sm a to sp or it es h ia n s 7 pe ri n op ol le n it es e la to id es 8 pi n u sp ol le n it es m in im u s 9 c er eb ro p ol le n ti es t h ie rg a rt ii 10 ve si ca sp or a f u sc u s alphabetical species list 1 baculatisporites sp. 5 bisaccate pollen 9 cerebropollenites thiergartii 6 chasmatosporites hians 3 deltoidospora spp. 2 lycopodiacidites rugulatus 7 perinopollenites elatoides 8 pinuspollenites minimus 4 rogalskaisporites cicatricosus 10 vesicaspora fuscus c: primulaelv 451.00 310.00 346753 346746 1 d el to id os p or a s pp . 2 b a cu la ti sp or it es s p. 3 s tr ia te lla j u ra ss ic a 4 m eg as po re s pp . 5 r et it ri le te s sp . 6 k ek ry p h a lo sp or a d is ti n ct a 7 k ra eu se lis p or it es r ei ss in ge ri i 8 is ch yo sp or it es v a ri eg a tu s 9 m a n u m ia d el co u rt ii 10 bi sa cc at e sp p. 11 c er eb ro p ol le n it es t h ie rg a rt ii 12 c or ol lin a t or os u s 13 c h a sm a to sp or it es h ia n s 14 pi n u sp ol le n it es m in im u s 15 s p h er ip ol le n it es s u b gr a n u la tu s 16 c er eb ro p ol le n it es m a cr ov er ru co su s 17 c h a sm a to sp or it es a p er tu s 18 c h a sm a to sp or it es m a jo r 19 n a n n oc er a to p si s sp . 20 n a n n oc er a to p si s se n ex 21 m a n co d in iu m s em it a b u la tu m 22 n a n n oc er a to p si s gr a ci lis 23 b ot ry oc oc cu s sp p. alphabetical species list 2 baculatisporites sp. 10 bisaccate spp. 23 botryococcus spp. 16 cerebropollenites macroverrucosus 11 cerebropollenites thiergartii 17 chasmatosporites apertus 13 chasmatosporites hians 18 chasmatosporites major 12 corollina torosus 1 deltoidospora spp. 8 ischyosporites variegatus 6 kekryphalospora distincta 7 kraeuselisporites reissingerii 21 mancodinium semitabulatum 9 manumia delcourtii 4 megaspore spp. 22 nannoceratopsis gracilis 20 nannoceratopsis senex 19 nannoceratopsis sp. 14 pinuspollenites minimus 5 retitriletes sp. 15 spheripollenites subgranulatus 3 striatella jurassica sy st em st ag e pa ly no lo gi ca l a ss em bl . z on es li th os tr at ig ra ph y l. p lie ns ba ch ia n g ul e h or n fm el is b je rg m b to ar ci an o st re ae lv f m n at h. f . m b sk æ v. m b lo w er ju ra ss ic (m ) sa m pl e he ig ht sa m pl e nu m be r interval not sampled 450 445 315 3101 5 rare few common abundant fig. 12. palynomorph distribution charts (for locations, see fig. 1). a, ostreaelv formation (nathorst fjeld member) at moskusoksekløft. b, gule horn formation (elis bjerg member) at harris fjeld. c, gule horn formation (elis bjerg member) and ostreaelv formation (nathorst fjeld (nath. f.) and skævdal (skæv.) members) at primulaelv. minimus and common botryococcus sp. the second sample (346745) was barren and thus does not appear on the distribution chart (fig. 12c). the uppermost sample is referred to assemblage zone 5 on the basis of abundant spheripollenites subgranulatus, cerebropollenites macroverrucosus and nannoceratopsis senex. suggested age. a late pliensbachian age is tentatively suggested for the lowermost sample (346746) based on a very poor assemblage in which only bisaccates and pinuspollenites minimus are common. an early toarcian age is suggested for the uppermost sample (346753) based on abundant spheripollenites subgranulatus, cerebropollenites macroverrucosus and nannoceratopsis senex. lepidopteriselv ten samples (139137–139146) were studied from this section and all of them are rich in palynomorphs (fig. 13, following page 744). these samples were collected by claus heinberg and tove birkelund in 1974, and they were thus not assigned to the recently-defined members. however, comparing their field notes with our sedimentological logs, it has been possible to assign the samples to the elis bjerg member. the samples 139237–139240 are rich in the spores deltoidospora and baculatisporites, the pollen pinuspollenites minimus and cerebropollenites thiergartii, bisaccate pollen and botryococcus sp. in sample 139141, the dinoflagellate cyst nannoceratopsis senex is very abundant and botryococcus sp. is rare. samples 139142–139146 are again rich in spores and pollen and botryococcus sp. whereas dinoflagellate cysts are rare. the palynological results allow us to suggest that the lowermost five samples (139137–139141) belong to assemblage zone 2. this is based on the common occurrence of cerebropollenites thiergartii and nannoceratopsis senex. the next five samples (139142–139146) are referred to assemblage zone 3 based on the presence of cerebropollenites thiergartii, pinuspollenites minimus and bisaccate pollen and the fact that dinoflagellate cysts are rare. 747 liaselv sy st em lo w er ju ra ss ic st ag e u pp er p lie ns ba ch ia n pa ly no lo gi ca l a ss em bl . z on es li th os tr at ig ra ph y g ul e h or n fo rm at io n el is b je rg m em be r (m ) 325 315 305 295 sa m pl e he ig ht 325.00 299.00 293.00 sa m pl e nu m be r 346665 346662 346667 1 b a cu la ti sp or it es s p. 2 d el to id os p or a s pp . 3 ly co p od ia ci d it es r u gu la tu s 4 r et it ri le te s se m im u ri s 5 n eo ra is tr ic k ia s p. 6 s tr ia te lla s ee b er ge n si s 7 d en so is p or it es s ca n ic u s 8 r et it ri le te s sp . 9 bi sa cc at e sp p. 10 pi n u sp ol le n it es m in im u s 11 pe ri n op ol le n it es e la to id es 12 c er eb ro p ol le n it es t h ie rg a rt ii 13 c h a sm a to sp or it es h ia n s 14 c er eb ro p ol le n it es m a cr ov er ru co su s 15 q u a d ra ec u lin a a n el la ef or m is 16 c h a sm a to sp or it es m a jo r 17 ve si ca sp or a f u sc u s 18 c or ol lin a t or os u s 19 n a n n oc er a to p si s se n ex 20 m en d ic od in iu m s p. 21 m a n co d in iu m s em it a b u la tu m 22 d in oc ys t sp p. 23 a cr ita rc h sp p. 24 b ot ry oc oc cu s sp p. ? alphabetical species list 23 acritarch spp. 1 baculatisporites sp. 9 bisaccate spp. 24 botryococcus spp. 14 cerebropollenites macroverrucosus 12 cerebropollenites thiergartii 13 chasmatosporites hians 16 chasmatosporites major 18 corollina torosus 2 deltoidospora spp. 7 densoisporites scanicus 22 dinocyst spp. 3 lycopodiacidites rugulatus 21 mancodinium semitabulatum 20 mendicodinium sp. 19 nannoceratopsis senex 5 neoraistrickia sp. 11 perinopollenites elatoides 10 pinuspollenites minimus 15 quadraeculina anellaeformis 4 retitriletes semimuris 8 retitriletes sp. 6 striatella seebergensis 17 vesicaspora fuscus 2 uncertain determination very rare rare few common abundant ? r fig. 14. palynomorph distribution chart for the gule horn formation (elis bjerg member) at liaselv (for location, see fig. 1). suggested age. a late pliensbachian age is proposed for the samples 139137–139146 based on the presence of cerebropollenites thiergartii, pinuspollenites minimus, nannoceratopsis species and mancodinium semitabulatum. liaselv three samples (346662, 346665, 346667) were analysed from the elis bjerg member (fig. 14); in general, preservation of the palynomorphs is very poor. all three samples are dominated by bisaccate pollen and botryococcus sp. and it is suggested that they belong to assemblage zone 2. suggested age. a late pliensbachian age has been suggested because of the abundance of pinuspollenites minimus and bisaccates. horsedal and the deltoidospora assemblage four samples (346696, 346700, 346701, 346703) were analysed from the horsedal member at horsedal (figs 15, 16). sample 346696 was barren. sample 346700, from a coal bed, yielded an assemblage composed predominantly of laevigate spores (pteridophyte spores) of the genus deltoidospora. such an assemblage has not been recorded before in samples from the neill klinter group at albuen or at any other locality in jameson land and scoresby land. the assemblage totally lacks microplankton. sample 346701 yielded a more diverse assemblage, but is dominated by bisaccate pollen and botryococcus sp. the uppermost sample, 346703, yielded only bisaccate pollen. it has not, based on the present material, been possible to place this assemblage within any of the assemblage zones defined above from the albuen succession. the assemblage is thus defined as a new assemblage, named the deltoidospora assemblage, which is presently only recognised at horsedal in the horsedal member of the ostreaelv formation. suggested age. this assemblage is not age specific; it could occur within any stage of the jurassic. palaeoenvironment. this assemblage is indicative of an enclosed swamp area (lagoon, pond, small lake) with a dense vegetation of ferns. ranunkeldal seven samples were analysed from this section (figs 5, 6). samples 341167–341170 were sampled in the uppermost part of the kap stewart group and samples 748 m si pb sand 815 820 346700 825 346703 346701 830 m horsedal w av e a nd s to rm -d om in at ed la go on d el to id os p or a a ss em bl ag e o st re ae lv f or m at io n h or se da l m em be r fig. 15. sedimentological log through part of the horsedal member (ostreaelv formation) in horsedal (for location, see fig. 1). sample numbers are indicated; arrows denote grain-size trends. for legend, see fig. 3. 341171–341173 are from the elis bjerg member of the gule horn formation. preservation of the palynomorphs from all these samples is very poor. however, the palynomorph assemblages are dominated by the laevigate spore deltoidospora sp., bisaccate pollen, chasmatosporites hians, cerebropollenites thiergartii and quadraeculina anellaeformis. botryococcus sp. is present in all samples, but is only abundant in the uppermost sample. in sample 341168, a dinoflagellate cyst has been found; it is similar to mendicodinium reticulatum, but shows some anomalous features. the occurrence of this dinoflagellate cyst suggests that the environment was influenced by brackish waters, at least for a short time. a single nannoceratopsis gracilis cyst was observed in sample 341171 from 306 m. in this sample, a tasmanites was found together with a poorly preserved foraminiferal inner-lining. in the uppermost sample, several leiofusa jurassica were recorded. the palynomorph assemblages from the kap stewart group in ranunkeldal are suggested to belong to a separate assemblage zone. the samples from the elis bjerg member are referred to assemblage zone 1. suggested age. a ?late sinemurian age is suggested for the kap stewart group samples because of the occurrence of the dinoflagellate comparable to mendicodinium reticulatum. this species has been found on bornholm, denmark in sediments of latest sinemurian and earliest pliensbachian age (batten et al. 1994; koppelhus & nielsen 1994). a pliensbachian age is suggested for the elis bjerg member samples because of the presence of nannoceratopsis gracilis. palaeoenvironment. a non-marine, freshwater environment is indicated for the kap stewart group samples, although the presence of a dinoflagellate cyst in sample 341168 suggests the influence of brackish water, albeit only temporarily. the samples from the elis bjerg member indicate marine influence. depositional environments and assemblage zones the palynological results presented above provide an additional dataset with which to constrain palaeoenvironmental and sequence stratigraphic interpretations. in the following section, the individual palynological assemblage zones, together with the deltoidospora assemblage, are discussed in relation to the sedimentological and stratigraphic data. assemblage zone 1: cerebropollenites thiergartii – pinuspollenites minimus – botryococcus this assemblage zone is characteristic of the sedimentary succession referred to sequence sq2 of dam & surlyk (1995, 1998) at qupaulakajik, albuen and 749 horsedal sy st em st ag e li th os tr at ig ra ph y o st re ae lv f or m at io n h or se da l m em be r pa ly no lo gi ca l a ss em bl ag e ? ? d el to id os p or a (m ) 825 815 sa m pl e he ig ht 828.00 823.00 813.00 sa m pl e nu m be r 346703 346701 346700 1 d el to id os p or a s pp . 2 b a cu la ti sp or it es s p. 3 c a la m os p or a t en er 4 r et it ri le te s sp . 5 n eo ra is tr ic k ia s p. 6 r et it ri le te s a u st ro cl a va to id es 7 ly co p od ia ci d it es r u gu la tu s 8 d en so is p or it es s ca n ic u s 9 c h a sm a to sp or it es m a jo r 10 pe ri n op ol le n it es e la to id es 11 c er eb ro p ol le n it es t h ie rg a rt ii 12 s p h er ip ol le n it es p si la tu s 13 c er eb ro p ol le n it es m a cr ov er ru co su s 14 c h a sm a to sp or it es h ia n s 15 pi n u sp ol le n it es m in im u s 16 bi sa cc at e sp p. 17 ve si ca sp or a f u sc u s 18 q u a d ra ec u lin a a n el la ef or m is 19 pa re od in ia h a lo sa 20 m en d ic od in iu m s p. 21 n a n n oc er a to p si s sp . 22 b ot ry oc oc cu s sp p. ? alphabetical species list 2 baculatisporites sp. 16 bisaccate spp. 22 botryococcus spp. 3 calamospora tener 13 cerebropollenites macroverrucosus 11 cerebropollenites thiergartii 14 chasmatosporites hians 9 chasmatosporites major 1 deltoidospora spp. 8 densoisporites scanicus 7 lycopodiacidites rugulatus 20 mendicodinium sp. 21 nannoceratopsis sp. 5 neoraistrickia sp. 19 pareodinia halosa 10 perinopollenites elatoides 15 pinuspollenites minimus 18 quadraeculina anellaeformis 6 retitriletes austroclavatoides 4 retitriletes sp. 12 spheripollenites psilatus 17 vesicaspora fuscus uncertain determination rare few common abundant ? fig. 16. palynomorph distribution chart for the horsedal member (ostreaelv formation) in horsedal (for location, see fig. 1). goniomyakløft, which consists of the upper part of the rævekløft formation and the lowermost part of the elis bjerg member (fig. 17). it is also present in the lower part of sequence sq3 at primulaelv, however, and in the lowermost part of the elis bjerg member in ranunkeldal, just above the kap stewart formation. the base of the assemblage zone thus coincides with the base of sequence sq1 (i.e. sb1) and the zone extends up into the lowermost part of sequence sq3 of dam & surlyk (1995, 1998). at the studied locations, the upper part of the rævekløft formation consists of cross-bedded, fossiliferous mediumto very coarse-grained sandstones (fig. 7), interpreted to represent fields of dunes or shoreface ridges on the shoreface (dam & surlyk 1995, 1998). in the elis bjerg member, the assemblage is present in subtidal sand sheet, shoreface and offshore transition deposits (figs 3a, 5). the rævekløft formation is capped by an important drowning surface that defines the base of the elis bjerg member (figs 7, 17; dam & surlyk 1995, 1998). the palynological assemblage is uniform, being dominated by bisaccate pollen and the freshwater to brackish alga botryococcus sp. spores include common deltoidospora and baculatisporites sp., and locally lycopodiacidites rugulatus. among the pollen, pinuspollenites minimus and cerebropollenites thiergartii are common. acritarchs are rare and dinoflagellate cysts are absent with the exception of one nannoceratopsis gracilis cyst and the undetermined dinoflagellate cyst (cf. mendicodinium reticulatum) from the ranunkeldal section. the assemblage probably reflects a vegetation with few fern species and several gymnosperm species. these grew close to a fresh or brackish water environment, where the botryococcus algae lived. when seen in the light of the strong marine indicators provided by the macrofossils, sedimentary structures and ichnology, it is suggested that this palynomorph assemblage is dominantly allochthonous, having been transported from a terrestrial to a shallow marine environment. similar palynomorph assemblages are known from other areas, for example in the uppermost sinemurian and lowermost pliensbachian of bornholm in the baltic sea (koppelhus & nielsen 1994). assemblage zone 2: nannoceratopsis–botryococcus the assemblage zone is characteristic of the middle part of the elis bjerg member at albuen, lepidopteriselv and liaselv in subtidal sand sheet and storm-dominated shoreface deposits, but also occurs in similar deposits in the lower part of the member at goniomyakløft and the upper part of the member at qupaulakajik (figs 2, 17). the assemblage zone is most characteristic of the lower part of sequence sq3 of dam & surlyk (1995, 1998), but is also locally present in the uppermost part of sequence sq2 (fig. 17). the top is placed below the transition from subtidal sand sheet deposits to tidal channel deposits of the elis bjerg member (fig. 3a). the palynomorph assemblage is dominated by the same spores, pollen and botryococcus sp. that characterise assemblage zone 1, but the incoming of nannoceratopsis senex, n. gracilis, parvocysta barbata and mancodinium semitabulatum together with limbicysta bjaerkei and a few more acritarchs indicates an increase in marine influence. in most of the samples, there are between three and seven different dinoflagellate cyst species and between one and five acritarch species. the marine interpretation of this palynomorph assemblage zone is in agreement with the sedimentological and ichnological data, which also indicate a shallow marine environment (dam & surlyk 1995, 1998). assemblage zone 3: chasmatosporites – cerebropollenites thiergartii – botryococcus the assemblage is characterised by rare dinoflagellate cysts such as mendicodinium reticulatum and nannoceratopsis spp. a few acritarchs are present, together with abundant pollen and botryococcus. the assemblage is present in the upper part of the elis bjerg member, but may extend into the lowermost part of the albuen member (fig. 17). assemblage zone 3 is dominantly present in stacked tidal channel and waveand storm-dominated shoreface deposits (fig. 3a; dam & surlyk 1995, 1998), whereas the sample from the albuen member was from heterolithic lower shoreface deposits. the base of the assemblage zone is placed just below the transition from subtidal sand sheet deposits to tidal channel deposits. at albuen, the top of the zone occurs just beneath the boundary between tidal channel deposits of the elis bjerg member and storm-dominated offshore transition deposits of the albuen member; this boundary has been interpreted as a coalesced sequence boundary and transgressive surface (dam & surlyk 1995, 1998). at astartekløft, however, the lowermost sample in the albuen member is also referred to assemblage zone 3 (fig. 10). 750 751 0 50 10 0 15 0 km 50 m h st h st h st h st h st t st t st sq 6 sq 5 sq 7 sq 4 sq 3 sq 2 sq 1 t re fjo rd b je rg qupaulakajik/ rævekløft albuen goniomyakløft astartekløft moskusoksekløft harris fjeld (n) nathorst fjeld dusén bjerg t id al c ha nn el s an d su bt id al s ho al s rhætelv horsedal halten terrace, norway lepidopteriselv/ liaselv primulaelv harris fjeld (s) a st ar te kl øf t/ h or se da l sk æ vd al m b fm sortehat ostreaelv gule horn a lb ue n el is b je rg r æ ve kl øf t t st t st t st h st t st t stls t n at ho rs t fj el d/ h ar ri s fj el d/ le pi do pt er is el v so ut h n or th sh or ef ac e t er m in al lo be la go on o ffs ho re t ra ns iti on fa ci es a ss oc ia tio n bo un da ry se qu en ce b ou nd ar y m aj or fl oo di ng s ur fa ce n ot ile r or t ilj e å re a ss em bl ag e z on e 4 a ss em bl ag e z on e 3 a ss em bl ag e z on e 2 a ss em bl ag e z on e 1 d el to id os p or a a ss em bl ag e a ss em bl ag e z on e 7 a ss em bl ag e z on e 6 a ss em bl ag e z on e 5 fi g. 1 7. n o rt h –s o u th c o rr el at io n p an el o f th e lo w er – lo w er m id d le j u ra ss ic n ei ll k lin te r g ro u p o f ja m es o n l an d , e as t g re en la n d s h o w in g th e d is tr ib u tio n o f th e p al yn o lo gi ca l a ss em b la ge z o n es r ep o rt ed h er e re la tiv e to th e m ai n s eq u en ce s tr at ig ra p h ic e le m en ts a n d th e d ep o si tio n al e n vi ro n m en ts . t h e co rr el at iv e fo rm at io n s o f th e h al te n t er ra ce , o ff sh o re n o rw ay , ar e sh o w n o n t h e ri gh th an d s id e o f th e fi gu re . n o te t h at p al yn o lo gi ca l d at a ar e sc ar ce n o rt h o f a st ar te kl ø ft . t h e d o tt ed v er tic al l in es i n d ic at e th e lo ca lit ie s o n w h ic h t h e se q u en ce st ra tig ra p h ic i n te rp re ta tio n o f d am & s u rl yk ( 19 98 ) is b as ed . a ss em b la ge z o n es 8 a n d 9 i n t h e so rt eh at f o rm at io n a re n o t in d ic at ed ; th e re ad er i s re fe rr ed t o k o p p el h u s & h an se n (2 00 3, t h is v o lu m e) . fi gu re m o d if ie d f ro m d am & s u rl yk ( 19 98 ). s q , se q u en ce ; ls t , lo w st an d s ys te m s tr ac t; t st , tr an sg re ss iv e sy st em s tr ac t; h st , h ig h st an d s ys te m s tr ac t. the shift from assemblage zone 2 to assemblage zone 3 records a change from a marine setting to an environment with a strongly terrestrial character, albeit with evidence of periodic marine influence. this floral/faunal change is compatible with the sedimentological record (dam & surlyk 1995, 1998), which indicates a change from subtidal sand sheet deposition in a shallow marine environment to a tidal channel environment, where a larger degree of terrestrial influence is to be expected. assemblage zone 4: bisaccates this palynomorph assemblage is very uniform, being dominated overwhelmingly by bisaccate pollen. the combined data from the albuen and astartekløft sections suggest that the assemblage characterises much of the albuen member of the gule horn formation. at albuen, the base of the assemblage zone occurs immediately above the coalesced sequence boundary and transgressive surface that separates the elis bjerg and albuen members (fig. 3b; sb4 of dam & surlyk 1995, 1998); at astartekløft, in contrast, assemblage zone 3 straddles this sequence boundary, assemblage zone 4 being recognised only in the uppermost levels of the albuen member (fig. 10). the upper levels of the albuen member proved inaccessible at albuen and were not sampled; the top of the zone is placed at the uppermost sample, some 12 m below the top of the member. at astartekløft, however, detailed sampling across the boundary between the albuen and astartekløft members demonstrated that the boundary between assemblage zones 4 and 5 coincides closely with this surface which is interpreted as a sequence boundary (figs 9, 10, 17; sb5 of dam & surlyk 1995, 1998). the albuen member is heterolithic, being composed of alternating mudstones and well-sorted fine-grained sandstones deposited in a storm-dominated lower shoreface environment (dam & surlyk 1995, 1998). a few coarse-grained pebbly sheets, moulded into large symmetrical ripples, and massive sandy mudstones deposited from debris flows are commonly interbedded with the heterolithic deposits. it is well-known that bisaccate pollen is commonly concentrated in distal marine settings, beyond the reach of other terrestrial palynomorphs. on this basis, this assemblage could therefore be interpreted to reflect an offshore marine environment, beyond the depositional range of other land-derived elements. the sedimentological data, however, testify to a lower shoreface environment (dam & surlyk 1998), and other land-derived elements should therefore be present. the absence of marine palynomorphs is also difficult to explain. as discussed earlier, such thin-walled forms may have been selectively destroyed by thermal effects caused by the intrusion of volcanics into the sediments. alternatively, the lack of marine palynomorphs could reflect partial isolation of the embayment resulting in the development of a fresh to brackish water environment; this could also explain the lack of tidal indicators in this member. assemblage zone 5: spheripollenites subgranulatus – cerebropollenites macroverrucosus – luehndea spinosa the base of assemblage zone 5 in the albuen section is characterised by the sudden incoming of spheripollenites subgranulatus and the reappearance of dinoflagellate cysts together with a more diverse pollen flora and the freshwater alga botryococcus sp. in sample 405466 at 259 m in the albuen section (figs 3b, 4a), luehndea spinosa appears for the first time together with common spherical dinocysts that are of unknown affinity, but have been recorded from spitsbergen, the danish subbasin and bornholm, denmark (bjærke 1980a; dybkjær 1991; koppelhus & nielsen 1994). at astartekløft, the lower zone boundary is placed just above the sequence boundary between the albuen and astartekløft members (figs 9, 17; sb5 of dam & surlyk 1995); the upwards extent of the zone is poorly constrained at this locality. at albuen, the upper boundary of the zone is placed approximately 10 m below the drowning surface that separates the heavily bioturbated shoreface sandstones of the nathorst fjeld member from bioturbated shelf deposits of the skævdal member (fig. 3c). at primulaelv, a single sample shows that this assemblage is also present just above the drowning surface (fig. 17). along neill klinter, the astartekløft member includes three facies associations, tidal channel, subtidal sand sheet and storm-dominated sandy shoal associations (fig. 3b). the tidal channel and subtidal sand sheet deposits are similar to those of the elis bjerg member, discussed above under assemblage zones 2 and 3. the storm-dominated sandy shoal deposits form a laterally continuous succession, composed of well-sorted fineto medium-grained sandstone beds (dam & surlyk 1995, 1998). the nathorst fjeld member forms a single coarsening-upwards succession consisting of alternating silty mudstones and 752 thin laminae of very fineto fine-grained sandstones, grading upwards into fineto coarse-grained sandstones. the sandstones are cross-bedded, wave ripple cross-laminated, hummocky cross-stratified and bioturbated. the coarsening-upwards succession reflects an increase in energy with time and is interpreted to record progressive shallowing from an offshore transition setting to a shoreface environment. the spheripollenites subgranulatus – cerebropollenites macroverrucosus – luehndea spinosa assemblage comprises a mixture of spores, pollen, a few dinoflagellate cyst species and acritarchs and the freshwater alga botryococcus. the assemblage zone indicates a brackish to marine environment with a large input of terrestrial material. this is in accordance with the sedimentological data indicating various environments in a marginal shallow marine setting. assemblage zone 6: perinopollenites elatoides assemblage zone 6, which is typical of the skævdal and the trefjord bjerg members (fig. 17) is characterised by the abundance of the pollen perinopollenites elatoides and the absence or scarcity of spheripollenites subgranulatus. overall, the palynomorph assemblage is dominated by different pollen species but dinoflagellate cysts are also significant, including dissiliodinium sp., phallocysta eumekes, pareodinia halosa, and kallospharidium sp.; acritarchs are also present. at albuen, in the south-eastern part of the basin, the lower boundary is placed some 10 m below the drowning surface that defines the top of the nathorst fjeld member (figs 3c, 17). the upper boundary of the assemblage zone at albuen is placed at 434 m in the upper trefjord bjerg member. succeeding samples, just below the boundary between the trefjord bjerg member and the sortehat formation, are referred to assemblage zone 7, which is characteristic of the lower sortehat formation (figs 3c, 17). at 437 m, between these two sampled levels, is an erosional surface that is draped by well-rounded quartzite pebbles up to 3 cm across; this surface is defined as a sequence boundary (sb7 of dam & surlyk 1998). the skævdal member consists of bioturbated muddy sandstones and deposition probably took place in a low-energy shelf environment (dam & surlyk 1995, 1998). primary physical structures only occur locally and include wave ripple cross-lamination, cross-lamination and cross-bedding. stratigraphic variations in the mud content suggests that the heavily bioturbated muddy sandstones were originally deposited as heteroliths (dam & surlyk 1995, 1998). the skævdal member is truncated by a prominent basinwide erosional unconformity, in places draped by a lag conglomerate. the unconformity marks a basinwide seawards shift in facies and is interpreted as a sequence boundary (sb6 of dam & surlyk 1995, 1998). at albuen, the sequence boundary is overlain by subtidal cross-bedded sandstones of the trefjord bjerg member deposited in an extensive subtidal dune field. the palynomorph assemblage zone 6 is indicative of deposition in a marine environment with a large input of terrigenous material. this is in accordance with the sedimentological data that indicate a shallow marine environment (dam & surlyk 1995, 1998). assemblage zone 7: botryococcus this assemblage is characterised by abundant botryococcus. in a few samples near the lower boundary of the zone, both botryococcus and dinoflagellate cysts occur in abundance, but the latter become rare upwards within the zone; the top of the zone is marked by the re-appearance of dinoflagellate cysts and the disappearance of botryococcus. the depositional environment of this assemblage zone is discussed in detail in an accompanying paper (koppelhus & hansen 2003, this volume). deltoidospora assemblage this assemblage is restricted to the horsedal member in the northern part of the basin (figs 15–17). this member is made up of minor coarsening-upwards successions, 1–6 m thick, deposited in wave-dominated beaches or delta systems that prograded into an extensive lagoonal environment (dam & surlyk 1995, 1998). the palynomorph assemblage is overwhelmingly dominated by laevigate spores (pteridophyte spores) of the genus deltoidospora and is suggestive of an enclosed swamp area (lagoon, pond, small lake) with a dense vegetation of ferns. this is in close agreement with the depositional environment suggested by sedimentary facies analysis (dam & surlyk 1995, 1998). discussion and conclusions seven palynological assemblage zones have been recognised in the rævekløft, gule horn and ostreaelv 753 formations of the neill klinter group (fig. 17). six of these are defined in this paper; the uppermost zone is defined by koppelhus & hansen (2003, this volume) in an accompanying paper as it is most characteristic of the overlying sortehat formation, the uppermost formation of the neill klinter group (koppelhus & hansen 2003, this volume). in addition, an assemblage termed the deltoidospora assemblage is defined here from the horsedal section. the seven palynological assemblage zones were all recognised primarily on the basis of data from the albuen section but additional data from other localities suggest that the zones may have a basinwide distribution (figs 17, 18). the palynological assemblages contain a diverse palynoflora, including 136 species. the assemblages indicate that the neill klinter group spans the early pliensbachian to early aalenian, without any major breaks in the stratigraphic record. the study 754 az7 az6 az5 az4 az5 az3 az3 az2 az1 sb7 sb6 sb5 sb4 sb3 sb2 sb1 a al en ia n to ar ci an la te p lie ns ba ch ia n so rt eh at f m o st re ae lv f m n ei l k lin te r g ro up g ul e h or n fm r æ ve kl øf t fm k s si n. ea rl y pl ie ns ba ch ia n sk æ vd al m b n at ho rs t fj el d m b a st ar te kl øf t m b el is b je rg m b a lb ue n m b tr ef jo rd bj er g m b ? az2 az5 az3 az2 az2 az1 az1 az2 az4 az1 az5 a lb ue n a st ar te kl øf t r æ ve kl øf t g on io m ya kl øf t le pi do pt er is el v li as el v r an un ke ld al h ar ri s fj el d pr im ul ae lv m os ku so ks ek lø ft az1 fig. 18. correlation diagram showing the stratigraphic distribution of the palynological assemblage zones (az1–7) at the studied locations (see fig. 1). note that the boundaries between assemblage zones 1 and 2 and between assemblage zones 3 and 4 are ‘diachronous’ relative to sequence boundaries sb3 and sb4 respectively. ks, kap stewart group; sin., sinemurian. 755 shows that the palynomorph flora of the neill klinter group is strongly influenced by the amount of organic matter derived from land plants and freshwater environments, yet the brackish and marine microplankton play a very important role in the interpretation of the environment and in establishing a palynostratigraphy for the group. comparing the sequence stratigraphic framework developed by dam & surlyk (1995, 1998) with the distribution of the palynological assemblages, it is clear that some of the major sequence stratigraphic and lithological boundaries are reflected by changes in the assemblages. assemblage zones 1–3 are characteristic of sequences sq1–3. figures 17 and 18 illustrate that the boundary between assemblages 1 and 2 is diachronous on a regional scale with respect to the important sequence stratigraphic surfaces (e.g. sb3). given that the sole difference between these two zones is the presboreal ammonite zones palyno-events in east greenland palyno-events on the mid-norwegian shelf opalinuma al en ia n pl ie ns ba ch ia n lo w er lo w er u pp er u pp er t oa rc ia n levesquei thouarsense variabilis bifrons falciferum tenuicostatum spinatum margaritatus davoei ibex jamesoni pareodinia halosa common botryococcus spp. acme callialasporites dampieri fad wallodinium laganum fad pareodinia halosa fad perinopollenites elatoides acme abundant cerebropollenites macroverrucosus luehndea spinosa fad cerebropollenites thiergartii becomes rare spheripollenites acme only bisaccate pollen cerebropollenites thiergartii dinoflagellate cysts parvocysta sp. fad mancodinium semitabulatum nannoceratopsis senex/gracilis fad abundant botryococcus spp. wallodinium laganum acme callialasporites dampieri fad increasing parvocysta sp. comparodinium sp. fad abundant perinopollenites elatoides nannoceratopsis gracilis/senex acme sphaeromorph clusters acme chasmatosporites sp. n. gracilis/senex acme sphaeromorph clusters acme chasmatosporites sp. n. gracilis/senex acme sphaeromorph clusters acme chasmatosporites sp. n. gracilis/senex acme spheripollenites acme abundant–common l. spinosa luehndea spinosa fad c. thiergartii present but becomes rare in younger sediments mancodinium semitabulatum n. senex/gracilis fad abundant–common botryococcus spp. common acritarchs fig. 19. diagram showing major palynoevents in the neill klinter group compared to the early jurassic palynoevents recorded from the halten terrace, mid-norwegian shelf (i. throndsen, personal communication 1996). fad, first appearance datum. 756 ence/absence of marine dinoflagellates, such diachroneity is not surprising. it can be attributed to variations in terrestrial input and the dominance of the freshwater plume laterally along the basin margin. the transition from assemblage zone 2 to assemblage zone 3 is marked by a decrease in dinoflagellate cysts which coincides with a gradual overall change from a dominance of subtidal sand sheet deposits to a dominance of tidal channel deposits in the upper part of the elis bjerg member (figs 3a, 17; dam & surlyk 1998). the strengthening of the terrestrial signal recorded by the change in the palynological assemblage is thus in accordance with the sedimentological record. at albuen, sequence boundary sb4 separates assemblage zones 3 and 4; at astartekløft, however, a sample from a few metres above the sequence boundary is referred to assemblage zone 3. this may be due to reworking of the uppermost sediments of the elis bjerg member in the underlying sequence. palynological data from the astartekløft section suggest that sb5 separates assemblage zones 4 and 5, whereas the transition from assemblage zone 5 to 6 appears to be diachronous (fig. 17). sequence boundary sb7 separates assemblage zones 6 and 7 (fig. 17). with respect to the sequence stratigraphic interpretation given by dam & surlyk (1995), an important conclusion of this study is that assemblage zone 7, characteristic of the lower part of the overlying sortehat formation (koppelhus & hansen 2003, this volume), also occurs in the uppermost few metres of the trefjord bjerg member at albuen, above a laterally persistent erosional surface with a conglomerate lag (figs 3c, 17). at other localities, a thin conglomerate layer separates the trefjord bjerg member and the sortehat formation. dam & surlyk (1995) interpreted the trefjord bjerg member – sortehat formation boundary as a coalesced sequence boundary and transgressive surface. however, the palynological data suggest that at albuen the sequence boundary should be placed beneath this boundary, at the conglomerate-draped erosional surface, and thus that the uppermost sandstones (c. 3 m thick) of the trefjord bjerg member represent a thin lowstand or transgressive package (figs 3c, 17). moreover, the palynological data suggest that the inferred major flooding surfaces in the sequence stratigraphic analysis (i.e. base albuen member and base sortehat formation), may in fact record partial isolation of the embayment from the seaway between norway and greenland causing a freshwater to brackish environment to develop. this would also explain the lack of tidal indicators in the albuen member. if this is the case, the albuen member does not represent transgressive and highstand deposits in a sequence stratigraphic sense, but rather a period of physical isolation of the basin from the sea. the palynostratigraphy has proven to be an important tool in confirming the high-resolution sequence stratigraphic correlation between east greenland and the halten terrace of the mid-norwegian shelf proposed by dam & surlyk (1995). this detailed palynological study should make it possible to correlate to other localities in the north atlantic and to the mid-norwegian area and the northern part of the north sea. the pattern of events seen in the east greenland material is mirrored by data from the halten terrace. thus, according to i. throndsen (personal communication 1996), the pliensbachian is characterised by common to abundant botryococcus; this is followed by the incoming of the dinoflagellate cysts nannoceratopsis gracilis, n. senex and mancodinium semitabulatum together with common to abundant cerebropollenites thiergartii pollen and abundant bisaccate pollen in the upper pliensbachian (fig. 19). the dinoflagellate cyst luehndea spinosa occurs in the uppermost pliensbachian and lowermost toarcian together with nannoceratopsis gracilis and n. senex followed by an acme of spheripollenites subgranulatus together with the sphaeromorph clusters. in the upper toarcian, the dinoflagellate cyst parvocysta complex appears together with abundant perinopollenites elatoides pollen and the first callialasporites dampieri appear together with the dinoflagellate cyst wallodinium in the lowermost aalenian. acknowledgements the work carried out in connection with this project by e.b.k. was supported by a grant from the efp-93 projects 1313/93-0010 and 0017. the work of g.d. has been supported by bp exploration operating company limited, london, and the carlsberg foundation. the manuscript was read by karen dybkjær, james b. riding, jon r. ineson and finn surlyk who offered many helpful suggestions. in particular, we thank ingar throndsen who placed unpublished data from the halten terrace at our disposal, henrik nøhr-hansen for patient assistance with the range charts and jon r. ineson for careful editing of the manuscript. references batten, d.j., koppelhus, e.b. & nielsen, l.h. 1994: uppermost triassic to middle jurassic palynofacies and palynomiscellanea in the danish basin and fennoscandian border 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(eds): petroleum geology of the north european margin, 303–315. london: graham & trotman for the norwegian petroleum society (npf). surlyk, f., hurst, j.m., piasecki, s., rolle, f., scholle, p.a., stemmerik, l. & thomsen, e. 1986: the permian of the western margin of the greenland sea – a future exploration target. in: halbouty, m.t. (ed.): future petroleum provinces of the world. american association of petroleum geologists memoir 40, 629–659. sykes, r.m. 1974: sedimentological studies in southern jameson land, east greenland. ii. offshore–estuarine regressive sequences in the neill klinter formation (pliensbachian– toarcian). bulletin of the geological society of denmark 23, 213–224. underhill, j.r. & partington, m.a. 1994: use of genetic sequence stratigraphy in defining and determining a regional tectonic control on the ‘mid-cimmerian unconformity’ – implications for north sea basin development and the global sea-level chart. in: weimer, p. & posamentier, h.w. (eds): siliciclastic sequence stratigraphy. recent developments and applications. american association of petroleum geologists memoir 58, 449–484. wall, d. 1965: microplankton, pollen, and spores from the lower jurassic in britain. micropaleontology 11, 151–190. ziegler, p.a. 1988: evolution of the arctic – north atlantic and the western tethys. american association of petroleum geologists memoir 43, 198 pp. 758 manuscript received 21 february 1997; revision accepted 1 june 2000. 759 appendix 1: list of all recorded palynomorph taxa miospores: anapiculatisporites sp. a. telephorus (pautsch) klaus 1960 annulispora folliculosa (rogalska) de jersey 1959 apiculatisporites parvispinosus (leschik) schulz 1963 a. sp. araucariacites australis cookson 1947 baculatisporites sp. (plate 1, fig. 12) b. wellmanii (couper) krutzsch 1959 bisaccates indeterminate (plate 2, fig. 9) callialasporites dampieri (balme) dev 1961 (plate 2, fig. 5) c. microvelatus schulz 1966 c. minus (tralau) guy 1971 (plate 2, fig. 3) c. sp. c. turbatus (balme) schulz 1967 calamospora tener (leschik) mädler 1964 camarozonozporites rudis (leschik) klaus 1960 c. sp. campenia sp. cerebropollenites macroverrucosus (thiergart) schulz 1967 (plate 3, fig. 4) c. sp. c. thiergartii schulz 1967 (plate 3, fig. 1) chasmatosporites apertus nilsson 1958 (plate 3, fig. 5) c. elegans nilsson 1958 c. hians nilsson 1958 c. major nilsson 1958 (plate 3, fig. 7) c. minor nilsson 1958 c. sp. chomotriletes minor (kedves) pocock 1970 c. sp. cibotiumspora jurienensis (balme) filatoff 1975 cingulizonates inequalis (mädler) lund 1977 conbaculatisporites mesozoicus klaus 1960 c. sp. corollina meyeriana (klaus) venkatachala & goczan 1964 c. sp. c. torosus (reissinger) cornet & traverse 1975 (plate 3, fig. 2) deltoidospora minor (couper) pocock 1970 d. spp. (plate 1, fig. 1) densoisporites scanicus tralau 1968 (plate 1, fig. 9) d. velatus weyland & krieger 1953 densosporites sp. d. variabilis (waltz) potonié & kremp 1956 eucommiidites major schulz 1967 e. troedsonii erdtman 1948 exesipollenites tumulus balme 1957 foraminisporis jurassicus schulz 1967 fungal spores iraquispora sp. ischyosporites crateris balme 1957 (plate 1, fig. 7) i. sp. i. variegatus (couper) schulz 1967 (plate 2, fig. 1) kekryphalospora distincta fenton & riding 1987 (plate 1, fig. 3) kraeuselisporites reissingeri (harris) morbey 1975 (plate 1, fig. 5) leptolepidites major l. sp. (plate 1, fig. 4) limbosporites lundbladii nilsson 1958 lycopodiacidites rugulatus (couper) schulz 1967 (plate 2, fig. 4) manumia delcourtii (pocock) dybkjær 1991 (plate 2, fig. 2) marattisporites scabratus couper 1958 megaspore fragments monosaccate pollen monosulcites punctatus orlowska-zwolinska 1966 (plate 2, fig. 6) murospora sp. neoraistrickia gristhorpensis (couper) tralau 1967 n. sp. n. taylori playford & dettmann 1965 ovalispollis ovalis krutzsch 1955 perinopollenites elatoides couper 1958 (plate 3, fig. 3) perinosporites thuringiacus schulz 1962 pinuspollenites minimus (couper) kemp 1970 (plate 2, fig. 8) polycingulatisporites circulus simoncsics & kedves 1961 p. triangularis (bolkhovitina) playford & dettmann 1965 quadraeculina anellaeformis malyavkina 1949 (plate 2, fig. 7) retitriletes austroclavatoides (cookson) döring et al. 1963 r. clavatoides (couper) döring et al. 1963 r. semimuris (danzé-corsin & laveine) mckellar 1974 r. sp. (plate 1, fig. 8) ricciisporites tuberculatus lundblad 1954 rogalskaisporites cicatricosus (rogalska) danzé-corsin & laveine 1963 (plate 1, fig. 10) sculptisporites aulosenensis (schulz) koppelhus 1992 sestrosporites pseudoalveolatus (couper) dettmann 1963 spheripollenites psilatus couper 1958 s. subgranulatus couper 1958 (plate 3, figs 6, 8) staplinisporites caminus (balme) pocock 1970 (plate 1, fig. 11) stereisporites antiquasporites (wilson & webster) dettmann 1963 stereisporites stereoides (potonié & venitz) h.d. pflug in: thomson & pflug 1953 s. sp. striatella jurassica mädler 1964 s. parva (li & shang) filatoff & price 1988 s. scania (nilsson) filatoff & price 1988 s. seebergensis mädler 1964 (plate 1, fig. 2) s. spp. striate pollen 760 taeniasporites rhaeticus schulz 1967 t. sp. taurocusporites verrucatus schulz 1967 (plate 1, fig. 6) tigrisporites microrugulatus schulz 1967 t. sp. todisporites major couper 1958 t. minor couper 1958 t. sp. triletes sp. tripartina variabilis malyavkina 1949 uvaesporites argenteaeformis (bolkhovitina) schulz 1967 u. sp. vesicaspora fuscus (pautsch) morbey 1975 vitreisporites pallidus (reissinger) nilsson 1958 v. sp. vittatina sp. zebrasporites interscriptus (thiergart) klaus 1960 phytoplankton: acritarch spp. baltisphaeridium sp. beaumontella caminuspina (wall) below 1987 b. delicata (wall) below 1987 b. sp. botryococcus spp. (plate 6, figs 6, 7) celyphus rallus batten 1985 (plate 7, fig. 1) c. spp. crassosphaera sp. cymatiosphaera sp. dapcodinium sp. dinocyst sp. (plate 4, figs 8, 9) dissiliodinium sp. (plate 6, figs 4, 5) kallosphaeridium sp. lecaniella foveata singh 1971 (plate 7, fig. 5) l. spp. leiofusa jurassica cookson & eisenack 1958 (plate 7, fig. 2) leiosphaeridia spp. limbicysta bjaerkei (smelror) macrae et al. 1996 (plate 6, figs 1–3) luehndea spinosa morgenroth 1970 (plate 5, fig. 5) mancodinium semitabulatum morgenroth 1970 (plate 5, figs 1, 2) m. sp. mendicodinium groenlandicum (pocock & sarjeant) davey 1979 m. reticulatum morgenroth 1970 (plate 5, figs 3, 4) m. sp. micrhystridium exilium wall 1965 m. fragile deflandre 1937 m. intromittum wall 1965 m. lymensis wall 1965 m. spp. m. stellatum deflandre 1942 m. wattonense wall 1965 nannoceratopsis ambonis (drugg) riding 1984 (plate 4, fig. 4) n. dictyoambonis riding 1984 n. gracilis alberti emend. van helden 1977 (plate 4, figs 1, 3) n. plegas drugg 1978 n. senex van helden 1977 (plate 4, fig. 2) n. sp. n. triangulata prauss 1987 n. triceras drugg 1978 pareodinia halosa (filatoff) prauss 1989 (plate 4, fig. 7) parvocysta barbata bjærke 1980 p. nasuta bjærke 1980 p. sp. phallocysta eumekes dörhöfer & davies 1980 (plate 4, figs 5, 6) p. elongata (beju) riding 1994 pterospermella spp. scriniocassis sp. susadinium scrofoides (dörhöfer & davies) below 1987 tasmanites sp. tetraporina compressa kondrat’ev 1963 (plate 7, fig. 6) valvaeodinium armatum morgenroth 1970 v. spp. veryhachium collectum wall 1965 (plate 7, fig. 3) v. formosum stockmans & williere 1960 (plate 7, fig. 4) v. irregulare de jekhowsky 1961 v. reductum (deunff) de jekhowsky 1961 v. sp. v. trispinosum (eisenack) deunff 1954 wallodinium laganum feist-burkhardt & monteil 1994 (plate 5, fig. 6) w. spp. others: foraminifera spp. foraminiferal linings (plate 7, fig. 7) haplophragmoides spp. miscellaneous 761 plates 1–7 762 plate 1 palynomorphs from the neill klinter group at the albuen section. the scale bar is 10 microns. for each of the illustrated specimens (plates 1–7), the england finder reference (efr) is given. fig. 1. deltoidospora sp. sample 405414-3, efr s291. fig. 2. striatella seebergensis. sample 405466-3, efr d34. fig. 3. kekryphalospora distincta. sample 405466-3, efr j383. fig. 4. leptolepidites sp. sample 405423-3, efr t40. fig. 5. kraeuselisporites reissingeri. sample 405419-3, efr d294. fig. 6. taurocusporites verrucatus. sample 405466-3, efr h273. fig. 7. ischyosporites crateris. sample 405449-3, efr w29. fig. 8. retitriletes sp. sample 405420-3, efr p50. fig. 9. densoisporites scanicus. sample 405449-3, efr p19. fig. 10. rogalskaisporites cicatricosus. sample 405419-3, efr l513. fig. 11. staplinisportes caminus. sample 405449-3, efr j462. fig. 12. baculatisporites sp. sample 405414-3, efr t353. 763 1 2 3 4 5 6 10 11 12 7 8 9 764 plate 2 palynomorphs from the neill klinter group at the albuen section. the scale bar is 10 microns. fig. 1. ischyosporites variegatus. sample 405464-3, efr g43. fig. 2. manumia delcourti. sample 405449-3, efr v244. fig. 3. callialasporites minus. sample 405449-3, efr j291. fig. 4. lycopodiacidites rugulatus. sample 405401-4, efr z263. fig. 5. callialasporites dampieri. sample 405449-3, efr j203. fig. 6. monosulcites punctatus. sample 405414-3, efr t31. fig. 7. quadraeculina anellaeformis. sample 405414-3, efr k431. fig. 8. pinuspollenites minimus. sample 405420-2, efr v212. fig. 9. bisaccate sp., full dimensions 90 x 70 microns. sample 405418-3, efr p462. 765 1 2 4 7 9 8 5 6 3 766 plate 3 palynomorphs from the neill klinter group at the albuen section. the scale bar is 10 microns. fig. 1. cerebropollenites thiergartii. sample 405401-4, efr u252. fig. 2. corollina torosus. sample 405466-3, efr d171. fig. 3. perinopollenites elatoides. sample 405453-3, efr g243. fig. 4. cerebropollenites macroverrucosus. sample 405454-3, efr y303. fig. 5. chasmatosporites apertus. sample 405401-4, efr y363. fig. 6. spheripollenites subgranulatus. sample 405459-3, efr n293. fig. 7. chasmatosporites major. sample 405423-3, efr f51. fig. 8. spheripollenites subgranulatus. sample 405459-3, efr n20. 767 1 4 7 8 5 6 2 3 768 plate 4 palynomorphs from the neill klinter group at the albuen section. the scale bar is 10 microns. fig. 1. nannoceratopsis gracilis. sample 405449-3, efr n20. fig. 2. n. senex. sample 405466-3, efr k211. fig. 3. n. gracilis. sample 405414-3, efr o404. fig. 4. n. ambonis. sample 405453-3, efr f383. fig. 5. phallocysta eumekes. sample 405426-3, efr v47. fig. 6. p. eumekes. sample 405459-3, efr y56. fig. 7. pareodinia halosa. sample 405454-3, efr v314. fig. 8. dinoflagellate sp. sample 405419-3, efr d40. fig. 9. dinoflagellate sp. sample 405453-3, efr e272. 769 1 4 7 5 8 2 3 6 9 770 plate 5 palynomorphs from the neill klinter group at the albuen section. the scale bar is 10 microns. fig. 1. mancodinium semitabulatum. sample 405466-3, efr h531. fig. 2. m. semitabulatum. sample 405411-3, efr h37. fig. 3. mendicodinium reticulatum. sample 405422-3, efr m292. fig. 4. m. reticulatum. sample 405420-3, efr f381. fig. 5. luehndea spinosa. sample 405466-3, efr f554. fig. 6. wallodinium laganum. sample 405452-3, efr f573 771 1 2 3 4 5 6 772 plate 6 palynomorphs from the neill klinter group at the albuen section. the scale bar is 10 microns. fig. 1. limbicysta bjaerkei. sample 405456-3, efr n244. fig. 2. l. bjaerkei. sample 405449-3, efr e221. fig. 3. l. bjaerkei. sample 405405-3, efr r484. fig. 4. dissiliodinium sp. sample 405454-3, efr p204. fig. 5. dissiliodinium sp. sample 405456-3, efr h304. fig. 6. botryococcus sp. sample 405401-4, efr x334. fig. 7. botryococcus sp. sample 405401-4, efr t242. 773 1 2 4 6 7 5 3 774 plate 7 palynomorphs from the neill klinter group at the albuen section. the scale bar is 10 microns; this scale bar is not applicable to figure 7. fig. 1. celyphus rallus. sample 405466-3, efr e21. fig. 2. leiofusa jurassica. sample 405414-3, efr o404. fig. 3. veryhachium collectum. sample 405414-3, efr o403. fig. 4. v. formosus. sample 405408-3, efr l353. fig. 5. lecaniella foveata. sample 405454-3, efr w401. fig. 6. tetraporina compressa. sample 405419-3, efr k504. fig. 7. foraminiferal lining, 132 microns in diameter. sample 405464-3, efr p292. 775 1 2 543 6 7 albuen (a) lo w er ju ra ss ic m a to ar ci an o st re ae lv f or m at io n g ul e h or n fo rm at io n sk æ vd al m em be r n at ho rs t fj el d m em be r a lb ue n m b a s el is b je rg m em be r tr ef jo rd b je rg m b l. p lie ns ba ch ia n u pp er p lie ns ba ch ia n 439.00 438.50 434.00 432.00 431.50 428.00 425.90 416.50 411.50 402.00 392.00 379.40 377.50 375.50 369.50 367.50 366.20 365.00 363.00 361.30 359.00 297.80 295.40 294.20 292.20 290.75 290.20 287.80 287.00 284.70 283.40 273.80 271.00 267.40 264.27 264.25 256.50 253.40 250.50 249.00 247.40 246.00 241.00 240.00 238.20 238.00 236.80 231.70 229.80 229.50 229.40 229.00 222.50 215.50 213.00 211.00 341248 397452 405449 341247 405450 405451 405452 405453 405454 405455 405456 405457 341243 405458 405459 341245 405460 341241 405462 405464 405466 405434 405433 405432 405431 405430 405429 341236 405428 405427 405426 405425 341235 405424 405423 405422 341234 405421 405420 341233 405419 405418 405417 405416 405414 405413 405411 405410 405408 341232 405406 405405 405404 405403 405402 405401 1 ro ga lsk ai sp or ite s ci ca tr ic os us 2 co nb ac ul at isp or ite s m es oz oi cu s 3 ly co po di ac id ite s ru gu la tu s 4 d el to id os po ra s pp . 5 re tit ril et es c la va to id es 6 re tit ril et es s em im ur is 7 re tit ril et es s p. 8 ba cu la tis po rit es s p. 9 ti gr isp or ite s m ic ro ru gu la tu s 10 fo ra m in isp or is ju ra ss ic us 11 kr ae us el isp or ite s re iss in ge ri 12 tr ip ar tin a va ria bi lis 13 st er ei sp or ite s st er eo id es 14 d en so isp or ite s sc an ic us 15 re tit ril et es a us tr oc la va to id es 16 to di sp or ite s m in or 17 ke kr yp ha lo sp or a di st in ct a 18 ci bo tiu m sp or ite s ju rie ne ns is 19 st ria te lla p ar va 20 to di sp or ite s m aj or 21 st ria te lla ju ra ss ic a 22 le pt ol ep id ite s sp . 23 z eb ra sp or ite s in te rs cr ip tu s 24 st er ei sp or ite s an tiq ua sp or ite s 25 u va es po rit es a rg en te ae fo rm is 26 tr ile te s sp . 27 st ria te lla s pp . 28 st ria te lla s ee be rg en sis 29 ch om ot ril et es s p. 30 m eg as po re fr ag m en ts 31 an ap ic ul at isp or ite s sp . 32 an ap ic ul at isp or ite s te le ph or us 33 n eo ra ist ric ki a sp . 34 ta ur oc us po rit es v er ru ca tu s 35 d en so isp or ite s ve la tu s 36 sc ul pt isp or ite s au lo se ne ns is 37 u va es po rit es s p. 38 to di sp or ite s sp . 39 m an um ia d el co ur tii 40 is ch yo sp or ite s va rie ga tu s 41 is ch yo sp or ite s cr at er is 42 st ap lin isp or ite s ca m in us 43 po lyc in gu la tis po rit es tr ia ng ul ar is 44 se st ro sp or ite s ps eu do al ve ol at us 45 po lyc in gu la tis po rit es c irc ul us 46 n eo ra ist ric ki a ta ylo rii 47 le pt ol ep id ite s m aj or 48 li m bo sp or ite s lu nd bl ad ii 49 m ur os po ra s p. 50 ch as m at os po rit es h ia ns 51 ce re br op ol le ni te s th ie rg ar tii 52 ve sic as po ra fu sc us 53 q ua dr ae cu lin a an el la ef or m is 54 pe rin op ol le ni te s el at oi de s 55 pi nu sp ol le ni te s m in im us 56 bi sa cc at e sp p. 57 ce re br op ol le ni te s m ac ro ve rr uc os us 58 ch as m at os po rit es m aj or 59 ch as m at os po rit es a pe rt us 60 co ro llin a to ro su s 61 m on os ul ci te s pu nc ta tu s 62 ch as m at os po rit es s p. 63 ch as m at os po rit es e le ga ns 64 m on os ac ca te s pp . 65 vi tr ei sp or ite s pa llid us 66 ca llia la sp or ite s da m pi er i 67 ca llia la sp or ite s sp . 68 ar au ca ria ci te s au st ra lis 69 eu co m m iid ite s tr oe ds on ii 70 ce re br op ol le ni te s sp . 71 sp he rip ol le ni te s su bg ra nu la tu s 72 ca m pe ni a sp . 73 ca llia la sp or ite s m ic ro ve la tu s 74 ca llia la sp or ite s m in us 75 ex es ip ol le ni te s tu m ul us 76 ca llia la sp or ite s tr ilo ba tu s 77 ca llia la sp or ite s tu rb at us 78 eu co m m iid ite s m aj or 79 ri cc iis po rit es tu be rc ul at us 80 ta en ia sp or ite s sp . 81 st ria te s pp . ? ? ? r r ? ? r r r r r r r r alphabetical species list 31 anapiculatisporites sp. 32 anapiculatisporites telephorus 68 araucariacites australis 8 baculatisporites sp. 56 bisaccate spp. 66 callialasporites dampieri 73 callialasporites microvelatus 74 callialasporites minus 67 callialasporites sp. 76 callialasporites trilobatus 77 callialasporites turbatus 72 campenia sp. 57 cerebropollenites macroverrucosus 70 cerebropollenites sp. 51 cerebropollenites thiergartii 59 chasmatosporites apertus 63 chasmatosporites elegans 50 chasmatosporites hians 58 chasmatosporites major 62 chasmatosporites sp. 29 chomotriletes sp. 18 cibotiumsporites jurienensis 2 conbaculatisporites mesozoicus 60 corollina torosus 4 deltoidospora spp. 14 densoisporites scanicus 35 densoisporites velatus 78 eucommiidites major 69 eucommiidites troedsonii 75 exesipollenites tumulus 10 foraminisporis jurassicus 41 ischyosporites crateris 40 ischyosporites variegatus 17 kekryphalospora distincta 11 kraeuselisporites reissingeri 47 leptolepidites major 22 leptolepidites sp. 48 limbosporites lundbladii 3 lycopodiacidites rugulatus 39 manumiadel courtii 30 megaspore fragments 64 monosaccate spp. 61 monosulcites punctatus 49 murospora sp. 33 neoraistrickia sp. 46 neoraistrickia taylorii 54 perinopollenites elatoides 55 pinuspollenites minimus 45 polycingulatisporites circulus 43 polycingulatisporites triangularis 53 quadraeculinaanellae formis 15 retitriletes austroclavatoides 5 retitriletes clavatoides 6 retitriletes semimuris 7 retitriletes sp. 79 ricciisporites tuberculatus 1 rogalskaisporites cicatricosus 36 sculptisporites aulosenensis 44 sestrosporites pseudoalveolatus 71 spheripollenites subgranulatus 42 staplinisporites caminus 24 stereisporites antiquasporites 13 stereisporites stereoides 81 striate sp. 21 striatella jurassica 19 striatella parva 28 striatella seebergensis 27 striatella sp. 80 taeniasporites sp. 34 taurocusporites verrucatus 9 tigrisporites microrugulatus 20 todisporites major 16 todisporites minor 38 todisporites sp. 26 triletes sp. 12 tripartina variabilis 25 uvaesporites argenteaeformis 37 uvaesporites sp. 52 vesicaspora fuscus 65 vitreisporites pallidus 23 zebrasporites interscriptus 400 380 360 300 280 260 240 220 420 interval not sampled sy st em st ag e pa ly no lo gi ca l a ss em bl . z on es li th os tr at ig ra ph y (m ) sa m pl e he ig ht sa m pl e nu m be r 7 6 5 4 3 2 1 uncertain determination very rare rare few common abundant ? r fig. 4a. terrestrial palynomorph distribution chart for the gule horn and ostreaelv formations at albuen (for location, see fig. 1). m, middle jurassic; a, aalenian; as, astartekløft member. albuen (b) 439.00 438.50 434.00 432.00 431.50 428.00 425.90 416.50 411.50 402.00 392.00 379.40 377.50 375.50 369.50 367.50 366.20 365.00 363.00 361.30 359.00 292.20 287.80 287.00 284.70 283.40 273.80 271.00 267.40 264.27 264.25 256.50 253.40 250.50 249.00 247.40 246.00 241.00 240.00 238.20 238.00 236.80 231.70 229.80 229.50 229.40 229.00 222.50 215.50 213.00 211.00 341248 397452 405449 341247 405450 405451 405452 405453 405454 405455 405456 405457 341243 405458 405459 341245 405460 341241 405462 405464 405466 405431 341236 405428 405427 405426 405425 341235 405424 405423 405422 341234 405421 405420 341233 405419 405418 405417 405416 405414 405413 405411 405410 405408 341232 405406 405405 405404 405403 405402 405401 1 n an no ce ra to ps is se ne x 2 n an no ce ra to ps is sp . 3 n an no ce ra to ps is gr ac ilis 4 m en di co di ni um r et ic ul at um 5 be au m on te lla c am in us pi na 6 n an no ce ra to ps is tr ia ng ul at a 7 n an no ce ra to ps is pl eg as 8 m an co di ni um s em ita bu la tu m 9 pa rv oc ys ta b ar ba ta 10 m en di co di ni um g ro en la nd ic um 11 be au m on te lla d el ic at a 12 m en di co di ni um s p. 13 lu eh nd ea s pi no sa 14 d in oc ys t sp . 15 va lva eo di ni um a rm at um 16 va lva eo di ni um s pp . 17 n an no ce ra to ps is am bo ni s 18 d iss ilio di ni um s p. 19 pa rv oc ys ta s p. 20 ph al lo cy st a eu m ek es 21 n an no ce ra to ps is tr ic er as 22 pa re od in ia h al os a 23 ka llo sp ha er id iu m s p. 24 sc rin io ca ss is sp . 25 ph al lo cy st a el on ga ta 26 w al lo di ni um s pp . 27 a cr ita rc h sp p. 28 le io fu sa ju ra ss ic a 29 le io sp ha er id ia s pp . 30 m ic rh ys tr id iu m ly m en sis 31 m ic rh ys tr id iu m s pp . 32 ve ry ha ch iu m fo rm os um 33 li m bi cy st a bj ae rk ei 34 ve ry ha ch iu m c ol le ct um 35 te tr ap or in a co m pr es sa 36 le ca ni el la fo ve at a 37 ve ry ha ch iu m r ed uc tu m 38 ve ry ha ch iu m ir re gu la re 39 m ic rh ys tr id iu m in tr om itt um 40 m ic rh ys tr id iu m w at to ne ns e 41 m ic rh ys tr id iu m fr ag ile 42 ve ry ha ch iu m tr isp in os um 43 m ic rh ys tr id iu m s te lla tu m 44 cy m at io sp ha er a sp . 45 be au m on te lla s p. 46 pt er os pe rm el la s pp . 47 bo tr yo co cc us s pp . 48 ta sm an ite s sp . 49 le ca ni el la s pp . 50 fu ng al s po re s 51 ce lyp hu s sp p. 52 h ap lo ph ra gm oi de s sp p. 53 m is ce lla ne ou s sp p. ? ? ? ? ? ? ? alphabetical species list 27 acritarch spp. 5 beaumontella caminuspina 11 beaumontella delicata 45 beaumontella sp. 47 botryococcus spp. 51 celyphus spp. 44 cymatiosphaera sp. 14 dinocyst sp. 18 dissiliodinium sp. 50 fungal spores 52 haplophragmoides spp. 23 kallosphaeridium sp. 36 lecaniella foveata 49 lecaniella spp. 28 leiofusa jurassica 29 leiosphaeridia spp. 33 limbicysta bjaerkei 13 luehndea spinosa 53 miscellaneous spp. 8 mancodinium semitabulatum 10 mendicodinium groenlandicum 4 mendicodinium reticulatum 12 mendicodinium sp. 41 micrhystridium fragile 39 micrhystridium intromittum 30 micrhystridium lymensis 31 micrhystridium spp. 43 micrhystridium stellatum 40 micrhystridium wattonense 17 nannoceratopsis ambonis 3 nannoceratopsis gracilis 7 nannoceratopsis plegas 1 nannoceratopsis senex 2 nannoceratopsis sp. 6 nannoceratopsis triangulata 21 nannoceratopsis triceras 22 pareodinia halosa 9 parvocysta barbata 19 parvocysta sp. 25 phallocysta elongata 20 phallocysta eumekes 46 pterospermella spp. 24 scriniocassis sp. 48 tasmanites sp. 35 tetraporina compressa 15 valvaeodinium armatum 16 valvaeodinium spp. 34 veryhachium collectum 32 veryhachium formosum 38 veryhachium irregulare 37 veryhachium reductum 42 veryhachium trispinosum 26 wallodinium spp. sy st em st ag e pa ly no lo gi ca l a ss em bl . z on es li th os tr at ig ra ph y (m ) sa m pl e he ig ht sa m pl e nu m be r interval not sampled 1 2 3 4 5 6 7 lo w er ju ra ss ic m a to ar ci an o st re ae lv f or m at io n g ul e h or n fo rm at io n sk æ vd al m em be r n at ho rs t fj el d m em be r a lb ue n m b a s el is b je rg m em be r tr ef jo rd b je rg m b l. p lie ns ba ch ia n u pp er p lie ns ba ch ia n 400 380 360 300 280 260 240 220 420 uncertain determination very rare rare few common abundant ? r fig. 4b. marine palynomorph distribution chart for the gule horn and ostreaelv formations at albuen (for location, see fig. 1). m, middle jurassic; a, aalenian; as, astartekløft member. lepidopteriselv lo w er ju ra ss ic u pp er p lie ns ba ch ia n g ul e h or n fo rm at io n el is b je rg m em be r 700.00 695.00 685.00 679.00 674.00 654.00 653.00 648.00 642.00 638.00 139146 139145 139144 139143 139142 139141 139140 139139 139138 139137 1 ba cu la tis po rit es s p. 2 d el to id os po ra s pp . 3 ca la m os po ra te ne r 4 to di sp or ite s m aj or 5 ro ga lsk ai sp or ite s ci ca tr ic os us 6 st ria te lla p ar va 7 re tit ril et es a us tr oc la va to id es 8 re tit ril et es c la va to id es 9 n eo ra ist ric ki a sp . 10 st er ei sp or ite s st er eo id es 11 re tit ril et es s em im ur is 12 kr ae us el isp or ite s re iss in ge rii 13 re tit ril et es s p. 14 ti gr isp or ite s m ic ro ru gu la tu s 15 ap ic ul at isp or ite s pa rv isp in os us 16 co nb ac ul at isp or ite s m es oz oi cu s 17 d en so isp or ite s sc an ic us 18 ly co po di ac id ite s ru gu la tu s 19 ira qu isp or a sp . 20 ti gr isp or ite s sp . 21 tr ip ar tin a va ria bi lis 22 ke kr yp ha lo sp or a di st in ct a 23 an ap ic ul at isp or ite s sp . 24 m ar at tii sp or ite s sc ab ra tu s 25 d el to id os po ra m in or 26 an nu lis po ra fo llic ul os a 27 ap ic ul at isp or ite s sp . 28 st ria te lla ju ra ss ic a 29 d en so sp or ite s sp . 30 d en so sp or ite s va ria bi lis 31 ca m ar oz on os po rit es r ud is 32 m ur os po ra s p. 33 q ua dr ae cu lin a an el la ef or m is 34 bi sa cc at e sp p. 35 pe rin op ol le ni te s el at oi de s 36 ch as m at os po rit es h ia ns 37 pi nu sp ol le ni te s m in im us 38 ch as m at os po rit es a pe rt us 39 ce re br op ol le ni te s th ie rg ar tii 40 co ro llin a to ro su s 41 m on os ul ci te s pu nc ta tu s 42 ce re br op ol le ni te s m ac ro ve rr uc os us 43 ve sic as po ra fu sc us 44 ch as m at os po rit es m aj or 45 ar au ca ria ci te s au st ra lis 46 sp he rip ol le ni te s ps ila tu s 47 ca llia la sp or ite s tu rb at us 48 ca llia la sp or ite s m in us 49 eu co m m iid ite s tr oe ds on ii 50 m on os ac ca te s pp . 51 co ro llin a sp p. 52 vi tta tin a sp . 53 m en di co di ni um r et ic ul at um 54 m an co di ni um s em ita bu la tu m 55 n an no ce ra to ps is se ne x 56 n an no ce ra to ps is tr ia ng ul at a 57 n an no ce ra to ps is gr ac ilis 58 n an no ce ra to ps is sp . 59 m ic rh ys tr id iu m in tr om itt um 60 ve ry ha ch iu m tr isu lc um 61 m ic rh ys tr id iu m fr ag ile 62 le ca ni el la s pp . 63 m ic rh ys tr id iu m ly m en sis 64 ve ry ha ch iu m s p. 65 bo tr yo co cc us s pp . 66 ta sm an ite s sp . ? r r ? r ? r ? alphabetical species list 23 anapiculatisporites sp. 26 annulispora folliculosa 15 apiculatisporites parvispinosus 27 apiculatisporites sp. 45 araucariacites australis 1 baculatisporites sp. 34 bisaccate spp. 65 botryococcus spp. 3 calamospora tener 48 callialasporites minus 47 callialasporites turbatus 31 camarozonosporites rudis 42 cerebropollenites macroverrucosus 39 cerebropollenites thiergartii 38 chasmatosporites apertus 36 chasmatosporites hians 44 chasmatosporites major 16 conbaculatisporites mesozoicus 51 corollina sp. 40 corollina torosus 25 deltoidospora minor 2 deltoidospora spp. 17 densoisporites scanicus 29 densosporites sp. 30 densosporites variabilis 49 eucommiidites troedsonii 19 iraquispora sp. 22 kekryphalospora distincta 12 kraeuselisporites reissingerii 62 lecaniella spp. 18 lycopodiacidites rugulatus 54 mancodinium semitabulatum 24 marattiisporites scabratus 53 mendicodinium reticulatum 61 micrhystridium fragile 59 micrhystridium intromittum 63 micrhystridium lymensis 50 monosaccate spp. 41 monosulcites punctatus 32 murospora sp. 57 nannoceratopsis gracilis 55 nannoceratopsis senex 58 nannoceratopsis sp. 56 nannoceratopsis triangulata 9 neoraistrickia sp. 35 perinopollenites elatoides 37 pinuspollenites minimus 33 quadraeculina anellaeformis 7 retitriletes austroclavatoides 8 retitriletes clavatoides 11 retitriletes semimuris 13 retitriletes sp. 5 rogalskaisporites cicatricosus 46 spheripollenites psilatus 10 stereisporites stereoides 28 striatella jurassica 6 striatella parva 66 tasmanites sp. 14 tigrisporites microrugulatus 20 tigrisporites sp. 4 todisporites major 21 tripartina variabilis 64 veryhachium sp. 60 veryhachium trisulcum 43 vesicaspora fuscus 52 vittatina sp. sy st em st ag e pa ly no lo gi ca l a ss em bl . z on es li th os tr at ig ra ph y (m ) sa m pl e he ig ht sa m pl e nu m be r 700 675 650 3 2 uncertain determination very rare rare few common abundant ? r fig. 13. palynomorph distribution chart for the gule horn formation (elis bjerg member) at lepidopteriselv (for location, see fig. 1). geological survey of denmark and greenland bulletin 26, 2012, 21-24 21 methane distribution in holocene marine sediments in the bornholm basin, southern scandinavia jørn bo jensen and rudolf endler the baltic sea is an ideal natural laboratory to study the methane cycle in the framework of diagenetic processes. in this paper we present preliminary geological mapping results from project baltic gas, a research project with the overall aim to contribute to the development of a scientific basis for long term sustainable use and protection of the baltic sea ecosystem. the baltic sea is a marginal sea with a strong permanent haline stratification, which leads to oxygen-poor bottom waters, and which is sometimes interrupted by oxygen-rich saltwater flowing in from the north sea. the history of the baltic sea has resulted in deposition of organic-rich holocene marine sediments that overlie glacial, late-glacial and early holocene organic-poor sediments. the aims of baltic gas, a project within the bonus-169 joint baltic sea research programme running from 2009 to 2011, were (1) to map the occurrence of free shallow gas in holocene sediments, (2) to quantify methane fluxes through the sediments and into the water column and the atmosphere, and (3) to investigate the processes and parameters governing methane generation and consumption. the contribution by the geological survey of denmark and greenland, reported here, was to map the thickness and structure of organicrich marine deposits and the distribution of gas-bearing sediments in co-operation with partners. the authors have also compiled acoustic data which were used to select sites for a comprehensive coring programme. the sediment cores were used for physical characterisation of the gas-bearing sediments and for biogeochemical analyses. these included measurements of the concentrations of methane, sulphide, sulphate, iron and other elements and compounds. here we present data from the bornholm basin, one of several key study areas (fig. 1). methods on shallow seismic profiles, the acoustic return signal is reduced in areas rich in gas-bearing sediments. the most pronounced reduction is seen when the frequency used during the seismic survey is near the resonance of the gas bubbles. their size controls the resonance frequency, and multi-frequency data from the project show maximum bubble resonance close to 4.2 khz, which indicates gas bubbles with a radius of 0.5–2 mm. many of the acoustic/gas relationships were established by anderson & hampton (1980a, b). acoustic data acquired during the project comprised swath bathymetry data, multibeam backscatter data, multifig. 1. map of the baltic sea region showing the location of the project baltic gas in the bornholm basin (red rectangle) and the location of other place names mentioned in the text. 1: eckernförder bucht, 2: mecklenburger bucht, 3: arkona basin. fig. 2. map of the bornholm basin showing the distribution of shallow seismic lines and deep faults. the location of the seismic profile of fig. 5 is shown in green. baltic sea sweden norway denmark 500 km poland north sea 1 2 3 10 km 15°e 55°10´n seismic line profile in fig. 5 fault © 2012 geus. geological survey of denmark and greenland bulletin 26, 21–24. open access: www.geus.dk/publications/bull 2222 frequency single beam data (5–100 khz), echo sounder data as well as high frequency seismic data. sediment acoustic work using a 5–100 khz signal was mainly carried out in areas known to be rich in gas at shallow depths in mecklenburger bucht, the arkona basin and the bornholm basin in the western baltic. data from parasound and an innomar sediment echo sounder were acquired simultaneously for all acoustic lines. extensive seismic data were acquired from the gas-rich part of the bornholm basin (figs 1, 2). the sediment acoustic records were used to select places for collection of water-column data and bottom sediments during the cruises. the simultaneous recording with different devices allowed comparison of the different responses to the occurrence of gas at shallow depths. the new seismic field data collected during project baltic gas was loaded onto a seismic work station and combined with seismic archive data from the same area. the seismic dataset was interpreted and combined with physical characteristics of the sediments plus additional seabed data to compile a map of the gas distribution in the baltic sea. multisensor core logging of 6–12 m long gravity cores were used for estimating the basic physical properties of sediments with and without gas bubbles. split cores were used for core description, sub-sampling and sedimentological analyses. gas distribution mapping in the bornholm basin as stated in the introduction, the main aim of the project was to produce a map of the seabed gas distribution in the fig. 3. model of syn-sedimentary infill in a half-graben. fig.4. map of the bornholm basin showing the bathymetry and deep faults (black stippled lines). the red arrows show inflow channels. 10 m glacial deposits bedrock sw ne free gas 5 km littorina sea mud ancylus lake clay yoldia sea clay baltic ice lake clay fig 5. seismic profile 2005-06222 obtained by an innomar sediment echo sounder (10 khz). the profile crosses the bornholm basin (for location see fig. 2). deposits from the baltic ice lake, the yoldia sea and the ancylus lake drape the glacial basin surface whereas mud from the littorina sea shows asymmetrical infill. acoustic blanking due to free gas occurs where the thickness of the organic-rich littorina mud exceeds 6–8 m. bottom of basin basin footwall uplift border fault rivers e d ge o f b as in sediment layers fau lt s car p >100 95–100 90–95 85–90 80–85 75–80 70–75 65–70 60–65 55–60 50–55 45–50 <45 depth (m) 10 km 23 baltic sea. such a map shows only the general pattern, so in order to understand the mechanism of gas production below the seabed it was necessary to make detailed studies in a few key areas. the bornholm basin was selected as one of these key areas, because it is well known for gas-rich sediments and because seismic data from previous surveys together with the new data made it possible to get a full coverage of the basin with limited supplementary work during the cruises (fig. 2). the bornholm basin is located north-east of the island of bornholm in an area that has been influenced by block faulting. it is possible that faulting continued into the holocene. the bornholm basin is bounded by major faults and has been interpreted as a half-graben (fig. 3; vejbæk 1985; wannäs & flodén 1994; sviridov et al. 1995) in which thick packages of late glacial and holocene sediments have been deposited. in spite of the post-half-graben sedimentation history, the present bathymetry clearly reflects the deeper structures (fig. 4). two inflow channels are found in the southeastern part of the basin. it is a characteristic feature of the basin that late glacial and early holocene clay deposits drape the glacial surface, whereas the marine holocene mud sediments form a wedgelike sediment body (fig. 5). this difference reflects different sedimentation mechanisms, from vertical settling of sediment particles to settling influenced by inflowing currents during the marine littorina sea stage. a map showing the thickness of the holocene marine mud based on the depth of the seabed and the bottom of the holocene marine mud is shown in fig. 6. a clear connection between the thickness of the mud and the down-faulted blocks is seen; the mud reaches a thickness of more than 12 m in the vicinity of the fault scarp of the half-graben, but only a few metres in the deepest, central part of the bornholm basin. acoustic blanking is seen in many seismic profiles. this is caused by scattering due to gas bubbles in the sediment. a gas distribution map has been compiled showing the depth from the seabed to the top of the acoustic gas front (fig. 7). experience tells us that a critical thickness of organic-rich holocene marine mud must be reached before free gas bubbles form. in the bornholm basin where water depths in the order of 90 m are found, acoustic blanking starts where the holocene organic-rich mud reaches a thickness of 6–8 m. the depth from the seabed to the top of the acoustic gas front is an important parameter in modelling methane fluxes. our study shows that in the bornholm basin the gas front is located less than 0.5 m below the seabed, in areas with the highest sedimentation rates of organic-rich mud. acoustic properties and physical characteristics of gas-charged sediments the data acquired from multi-sensor core logging and sedimentological analyses were used for geo-acoustic models and interpretation of the seismo-acoustic records. the physical properties and the geo-acoustical data were used to investigate the influence of gas bubbles on the acoustic properties and the strength of the muddy sediments. the acoustic properties such as sound velocity and attenuation are strongly influenced by gas bubbles in the sediments, as illustrated in the sediment echo sounder and seismic records (fig. 5). the behaviour of acoustic signals is very complex and controlled by environmental parameters including pressure and temperature, the sound frequency and the physical properties of the different sedimentary components (solid grains, water and gas bubbles). the physical properties determined on sediment cores or samples from sediment fig. 6. map of the bormholm basin showing the thickness of holocene marine mud. deeper structures are represented by the faults (black stippled lines). 4.5–5.0 4.0–4.5 3.5–4.0 3.0–3.5 2.5–3.0 2.0–2.5 1.5–2.0 1.0–1.5 0.5–1.0 0.0–0.5 depth (m) >5 10 km fig. 7. mapped areas with acoustic blanks caused by scattering due to gas bubbles in the sediment. the depth from the seabed to the top of the acoustic gas front is shown. >12 10–12 8–10 6–8 4–6 2–4 0–2 thickness (m) 10 km 2424 samples (wet bulk density, porosity, gravimetric bulk water content, loss on ignition and vane shear strength) showed no significant differences between gas-charged and gas-free mud. the reason may be that gas bubbles only constitute a small part of the sediment and hence do not change the bulk properties significantly. similar results were reported by wilkens & richardson (1998) for eckernförder bucht where the gas volume ranged from 0 to 2%, with bubble diameters of 0.6–1 mm. concluding remarks ma pping of the methane distribution in the baltic sea surface sediments and sediment analyses have led to a number of conclusions: 1. acoustic blanking in seismic profi les is caused by scattering due to gas bubbles in the sediment. th e bubble size controls the resonance frequency, and the results indicate a maximum bubble resonance close to 4.2 khz, corresponding to bubble diameters in the order of 1–4 mm. 2. accumulations of near-surface gas in the baltic sea in general are restricted to near-shore archipelagos and geologically controlled sediment traps with high sedimentation rates. 3. a direct link between the occurrence of near-surface gas and a minimum thickness of holocene marine organic-rich sediments is seen. a characteristic feature for the bornholm basin is that acoustic blanks are seen where the organic-rich mud reaches a thickness of more than 6–8 m. 4. th e acoustic properties of gas-free and gas-charged sediments are very diff erent, as demonstrated by soundvelocity changes and attenuation in the seismic records. th e behaviour of acoustic signals is extremely complex and infl uenced by environmental parameters such as pressure and temperature. 5. comparisons of gas-charged and gas-free mud show no signifi cant diff erence in bulk properties of various basic physical parameters determined in the sediment. th is may refl ect that gas bubbles only constitute 0–5% of the holocene muddy sediments. acknowledgement project baltic gas was funded by the baltic organisations network (bonus). references anderson, a.l. & hampton, l.d. 1980a: acoustics of gas-bearing sediments i. background. journal of the acoustical society of america 67, 1865–1889. anderson, a.l. & hampton, l.d. 1980b: acoustics of gas-bearing sediments ii. measurements and models. journal of the acoustical society of america 67, 1890–1903. sviridov, n.i., frandsen, j.v., larsen, t.h., friis-christensen, v., madsen, k.e. & lykke-andersen, h. 1995: the geology of bornholm basin. aarhus geoscience 5, 15–35. vejbæk, o.v. 1985: seismic stratigraphy and tectonics of sedimentary basins around bornholm southern baltic. danmarks geologiske undersøgelse, serie a 8, 30 pp. wannäs, k.o. & flodén, t. 1994: tectonic framework of the hanö bay area, southern baltic sea. technical report: svensk kärnbränslehante ring ab, 1994, 50 pp., http://www.skb.se/upload/publications/pdf/ tr94-09webb.pdf wilkens, r.h. & richardson, m.d. 1998: the influence of gas bubbles on sediment acoustic properties: in situ, laboratory, and theoretical results from eckernförde bay, baltic sea. continental shelf research 18, 1859–1892. authors’ addresses j.b.j., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: jbj@geus.dk r.e., the leibniz institute for baltic sea research, seestrasse 15, warnemünde, d-18119 rostock, germany. geological survey of denmark and greenland bulletin 28, 2013, 29-32 29 geological map of denmark 1:50 000 – map sheet mors, nw denmark stig a. schack pedersen, peter roll jakobsen, lisbeth tougaard and peter gravesen danish geological maps of deposits occurring at the terrain surface are published under the name of geological map of denmark 1:50  000 and are based on geological field mapping at 1:25 000. most of the published maps follow the map sheet division shown in fig. 1. however, in some instances it is appropriate to publish geological maps covering a regional unit, such as an island. hence, the geological map of mors appears as the 1:50 000 map sheet mors, which covers parts of map sheets 1116 i, 1116 ii and 1116 iii (figs 1, 2; pedersen & jakobsen 2012). mors shows spectacular examples of glaciotectonic structures that are beautifully exposed in coastal cliff sections. it also has a unique geological history, which is briefly described in this paper. geological features of map sheet mors the geological map of mors shows deposits that are present at the terrain surface; they are mainly non-lithified quaternary deposits. the mapping was carried out using 1 m long hand augers. the soil and its underlying unit were tested at c. 100 m intervals with the auger, which collects a sample in a groove at its tip. the auger samples are classified in the field and allotted a symbol on a 1:25 000 field map. during more than 100 years of systematic mapping, about 50 symbols have been established which are used by the mapping geologist. the aim of the work is to map the boundaries between various units shown as polygons on the maps. the testing distance of 100 m provides a semi-statistical documentation of the recognised polygons. occasional outcrops of pre-quaternary deposits are classified with letter symbols. on the geological map of mors, the pre-quaternary geology is presented as an inset with structural contour lines showing the elevation of the prequaternary surface and the geological units that occur below the quaternary deposits (figs 2, 3). the map is mainly based on information from the jupiter well data base, but detailed geophysical mapping of mors (jørgensen et al. 2005) supported the interpretation of the structural contours. the pre-quaternary geology the pre-quaternary geological features on mors are influenced by the erslev structure, a salt diapir in the central part of the island (larsen & baumann 1982), and by glaciotectonic deformation (gry 1940; pedersen 2000). this is illustrated by three cross-sections documenting the relationship between the pre-quaternary and the quaternary geology (fig. 2). the diameter of the circular salt diapir is 5–6 km, © 2013 geus. geological survey of denmark and greenland bulletin 28, 29–32. open access: www.geus.dk/publications/bull published map sheets mapped areas unmapped areas 1318 i 1318 ii1318 iii 1217 i 1317 iv 1317 i 1216 ii 1316 iii 1315 iii 1217 ii 1315 i1315 iv1215 i 1216 i 1316 iv 1415 iv 1317 ii1317 iii 1812 iii 1812 iv 1417 iv 1416 ii 1215 ii 1214 i 1514 iv1414 i1314 i 1415 iii1315 ii 1213 iii 1513 iv1413 i1413 iv 1514 ii1414 ii1314 ii1214 ii 1314 iv 1214 iii 1214 iv 1514 iii1414 iii1314 iii 1213 iv 1215 iii 1313 i1313 iv 1212 iii 1413 ii1413 iii1313 ii1313 iii1213 ii 1213 i 1513 iii 1412 i1412 iv1312 i1312 iv1212 i1212 iv 1511 iii1411 ii1411 iii1311 ii 1511 iv1411 i 1411 iv1311 i1311 iv1211 i 1512 iii1412 ii1412 iii1312 ii1312 iii1212 ii 1211 iv 1112 iii 1112 i1112 iv 1112 ii 1111 i 1013 ii 1113 i1113 iv 1013 i 1113 i1113 iv 1114 ii1114 iii 1115 ii1115 iii 1115 iv 1115 i 1215 iv 1116 iii 1116 ii 1216 iii 1116 iv 1216 iv1116 i 1117 iii 1114 i1114 iv 1512 i1512 iv 1513 ii 1513 iv 1514 i 1217 iii1117 ii 1511 i 1512 ii 1510 iv mors 10°e 50 km 57°n jylland fig. 1. map of denmark showing the sheet divisions of the 1:50 000 geological map series. the red frames show map sheets published under the name of geological map of denmark 1:50 000. the map sheets are numbered according to the 1:50  000 topographic maps implemented by the former danish geodetic institute (now the danish geodata agency) in 1953. the map sheet division is based on the utm system and uses the european datum 1950. the survey decided to retain this map sheet division for the geological mapping, although the national survey and cadastre (now the danish geodata agency) in 2003 started to use the european terrestrial reference system 1989. 3030 and the diapir rises vertically from the base of the permian salt at a depth of 5–6 km. the top of the salt in the diapir is now found at depths of 600–700 m. a circular dome of chalk caps the salt diapir. maastrichtian chalk in its centre was formerly quarried in chalk pits, exposing the cretaceous–tertiary boundary. the chalk is overlain by bryozoan cherty limestone that forms an aureole around the top of the diapir, and paleocene clay occurs along the steeply dipping flanks. the uplift of the salt diapir continued in the paleocene and resulted in an increased thickness of plastic clay away from the diapir centre. in the surrounding marginal depression, eocene clayey diatomite occurs interbedded with volcanic ash layers. this unit, locally called ‘moler’, crops out in an old clay pit on the northern flank of the diapir (pedersen 2000). the moler is the most characteristic pre-quaternary unit in the glaciotectonic complexes that occur on northern mors. the diatomite is exploited and used for light-weight granulates. on northern mors, glaciotectonic folds are well exposed in the two largest clay pits. these folds and the surrounding hilly terrain were formed by superimposed glaciotectonic deformation (pedersen 2000). the most impressive glaciotectonic complex with up-thrust sheets of moler is the hanklit complex (figs 4, 5), where a thrust sheet is exposed that was displaced c. 300 m towards the foreland in the south. parallel ridges trend from hanklit towards the west, and also here the moler is the target of diatomite exploitation in pits following the crests of the anticlines. fig. 2. strongly reduced and modified version of map sheet mors with three cross-sections: a n–s section (upper), an e–w section (middle) and a ne–sw section of the northern part of mors (lower). the inset shows the depression on central mors which coincides with the depression in the top of the chalk covering the salt diapir (fig. 3). fig. 4(( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( (( prækvartæroverfladens geologi / geology of the pre-quaternary surface nedre miocæn, ler og sand lower miocene, clay and sand oligocæn, ler oligocene, clay eocæn, moler (diatomit med askelag) eocene, clayey diatomite with ash layers palæocæn, plastisk ler paleocene, clay danien, kalk danian, limestone øvre kridt, skrivekridt upper cretaceous, chalk prækvartæroverfladens højdekurver i meter contours in meters of the pre-quaternary surface normalforkastning normal fault glacialtektonisk forkastning glacial tectonic thrust fault 0 5 km 0 1 2 km kurveinterval: 5 m / contours interval: 5 m by, anlæg og fyld town, constructions and dumps råstofgrav raw material pit nedlagt råstofgrav raw material pit, abandoned postglaciale aflejringer postglacial deposits senglaciale aflejringer lateglacial deposits glaciale aflejringer glacial deposits ferskvandsler freshwater clay ferskvandssand freshwater sand ferskvandsgrus freshwater gravel ferskvandsgytje freshwater gyttja ferskvandstørv freshwater peat okker og myremalm ochre and iron ore saltvandsler marine clay saltvandssand marine sand saltvandsgrus marine gravel saltvands skalgrus marine shell gravel vekslende saltvandsaflejringer alternating marine deposits flyvesand aeolian sand vekslende tynde saltvandslag, marsk alternating marsh deposits smeltevandsler meltwater clay smeltevandssand meltwater sand smeltevandsgrus meltwater gravel moræneler till, clayey morænesand till, sandy kalkmorænesand chalky sandy till vekslende tynde smeltevandslag alternating thin meltwater deposits flodslette ler og finsand outwash clay and fine sand terrassesand terrace sand terrassegrus terrace gravel 0 -25 -50 -75 e f 6310000 56°55'0 56°50'0 6300000 56°45'0 6290000 56°40'0 6280000 56°55' 6300000 56°50' 6290000 6280000 56°45' 56°40' 8°30' 470000 8°35' 8°40' 480000 8°45' 490000 8°55' gullerup stærhøj mosebjerg klovbakkerne sundby stengrund hanklit flade klit skærbæk klint harhøj ejerslev lyng ejerslev havn ejerslev klint nees øre holmene glomstrup vig karby odde trædemark odde ørndrup hage stokkær odde flejskær hage skallerup hage thisted bredning limfjorden visby bredning dragstrup vig vi ls un d kås bredning sa ll in g su nd livø bredning assels hage søndervig lindholm stenklipperne a g e r ø dover mølle grund karby vig hage gudnæs skyum øre å bæ k 5 1 6 5 8 8 8 5 7 5 9 8 5 2 1 3 8 2 25 7 17 16 1 5 7 6 7 1 9 8 4 2 9 18 3 7 1 5 5 11 1 9 5 14 12 11 12 10 11 10 3 18 16 1 1 14 3 1 1 11 1 7 1 1 1 1 ejerslev røn erslev kærgl. jølby frøslev vang sindbjerg a b e f c d flade sejerslev ejerslev 0 -25 -50 -75 a bsønderby trehøj vilscentrum frøslevvang ersleverslev kær hanklit 0 0 0 0 0 0 0 -25 -25 -25 -25 0 0 -25 -25 -50 -25 -25 -25 -25 -25 -25 -25 -25 -25 -25 -100 -75 -75 -75 -50 -25 -25 -50 -50 -25 -25 -50 +25 -75 -50 -50 -25 -25 +25 +25 0 0 0 0 -25 0 0 0 0 0 -50 -50 -50 -50 -75 -50 0 -25 -50 dc visby bredning damsgård mollerup frøslevvang nykøbing mors thisted saltdome erslev saltdiapir hannæs fe gge su nd sø bugt salgerhøj ( (( ( fig. 4 31 between the glaciotectonic complexes on northern mors and the northern flank of the salt diapir, red and green plastic clay of the røsnæs ler formation occurs together with dark brown, micaceous, oligocene to early miocene clay. local outcrops of the oligocene–miocene clay (the brejning and viborg formations) are known in the hills along the strait west of mors. south of the erslev structure, miocene heterolithic deposits with mica-rich clay and sand are known from wells and were recognised during the systematic mapping of southern mors. the glaciodynamic geology the oldest quaternary deposits on mors comprise glaciolacustrine clay that is referred to the elsterian glaciation exposed in coastal cliffs on south-western mors. on northern mors, alternating beds of till and glaciofluvial sand are exposed along the west coast, and similar elsterian deposits are known from wells penetrating the buried tunnel valleys. a coastal cliff along the west side of northern mors exposes 8 m thick sandy till, rich in chalk, which was deposited by a norwegian ice advance during the saalian glaciation (about 300 000 years bp). characteristic erratic blocks of chalk and flint, originating from erosion of the till, are abundant along the shore (pedersen et al. 2012). this till and the underlying glaciofluvial sand and gravel are well known from diatomaceous clay pits on north-eastern mors. glaciolacustrine clay is widespread over large parts of northern and north-western mors. the clay was deposited during the middle weichselian prior to the norwegian ice advance. the depocentre was located in the depression between the thisted salt dome north of mors and the erslev salt diapir. both these areas formed elevated terrains, whereas the remaining part of mors constituted lowlands characterised by lake-filled depressions with glaciofluvial sand and gravel. in the area around hanklit, glaciolacustrine clay forms an important element in the glaciotectonic complex (figs 4, 5). the clay is thrust up into sheets, forming a terrain with e–w-trending parallel ridges. the arc-formed complex was created by the norwegian ice advance about 27 000 years bp. a lodgement till rich in indicator boulders such as larvikite and rhombohedral porphyries from the oslo region was deposited during this advance. fig. 3. bedrock map of mors (enlarged version of inset in fig. 2). the structural contour lines at 25 m intervals show the elevation of the pre-quaternary surface. the dome-like structure on central mors is made of chalk that covers the top of the erslev salt diapir. the boundaries between glaciotectonic complexes and their foreland on northern mors are also shown. fig. 4. part of the map sheet showing the hilly terrain at flade klit. the hanklit glaciotectonic complex is marked by a landscape dominated by elongated hills. the fur formation has been thrust-faulted up into parallel ridges in the complex. farther to the south, the thrust sheets mainly consist of glaciolacustrine clay and glaciofluvial sand. contour interval 5 m. for location see fig. 2. 26 26 26 581 581 581 26 26 26 581 581 581 581 26 mosebjerg klovbakkerne skærbæk klint stokkær odde flejskær hage skallerup hage vilsund øst vilsundbroen under bakken 8 5 1 3 18 16 1 1 14 3 7 erslev kærgl. jølby gullerup stærhøj sundby stengrund hanklit flade klit 1 b e f salgerhøj 1 km outwash clay and sand fur formation holocene deposits late glacial deposits glacial deposits tertiary deposits freshwater clay freshwater gyttja peat marine sand marine gravel glaciolacustrine clay glaciofluvial sand glaciofluvial gravel clayey till 56°50´n 8°48´e –50 –50 normal fault glaciotectonic thrust fault contours of the pre-quaternary surface (m) thisted salt dome erslev salt diapir 0 0 0 0 –25 –25 0 –25 –50 –25 –25 –75 –75 +25 –75 0 0 –50 –50 –50 –50 –25 5 km lower miocene clay and sand oligocene clay eocene diatomite & ash layers paleocene clay danian limestone upper cretaceous chalk 3232 after the norwegian ice melted back, mors was characterised by parallel hills intersected by lake-filled valleys with glaciolacustrine clay deposits as patches scattered over the landscape. a few thousand years later the advancing swedish ice remodelled the landscape. on northern mors, thrusting and folding from the north-east superimposed the architecture of the glaciomorphological landscape. when the ice reached south-western mors it stopped briefly and formed a stationary line, and its meltwater created a cone-shaped outwash plain. at its apex, boulders, stones and gravel were deposited, which are now quarried in gravel pits. the postglacial geology during the late glacial and holocene periods, mors was affected by three significant events: (1) meltwater erosion during the deglaciation and formation of a kettle-hole landscape caused by melting of dead ice, (2) marine transgression accompanied by sedimentation in straits and fjords after the ice age, and (3) glacio-isostatic rebound. during the atlantic sea-level highstand, coastal plains formed in front of the fossil coastal cliffs. a number of fossil straits and fjords with marine gyttja and marine sand occur on mors. the most significant fossil fjord is an e–w-trending depression located above the central part of the salt diapir on central mors. marine heterolithic deposits are characterised by shells of cardium, and mounds of oyster (ostrea edulis) beds formed where strong currents cut narrow gaps in the former fjord. a number of distinctive fossil coastal cliffs occur, particularly on northern mors where raised shorelines are found up to c. 5 m above the present sea level. the raised beaches have occasionally dammed coastal lakes. beach ridges occur along the coastline on southern mors. several of these contribute to the closure of small fjords and coastal lakes with peat accumulation. in the past, the coastal plains were larger, but are now subjected to increased erosion along the coastline of mors. references gry, h. 1940: de istektoniske forhold i moleret. med bemærkninger om vore dislocerede klinters dannelse og om den negative askeserie. meddelelser fra dansk geologisk forening 9, 586–627. jørgensen, f., sandersen, p.b.e., auken, e., lykke-andersen, h. & sørensen, k. 2005: contributions to the geological mapping of mors, denmark – a study based on a large-scale tem survey. bulletin of the geological society of denmark 52, 53–75. larsen, g. & baumann, j. 1982: træk af mors salthorstens udvikling. dansk geologisk forening, årsskrift for 1981, 151–155. pedersen, g.k et al. 2012: molerområdets geologi – sedimenter, fossiler, askelag og glacialtektonik. geologisk tidsskrift 2011, 41–135. pedersen, s.a.s. 2000: superimposed deformation in glaciotectonics. bulletin of the geological society of denmark 46, 125–144. pedersen, s.a.s. & jakobsen, p.r. 2012: geological map of denmark, 1:50 000, mors. copenhagen: geological survey of denmark and greenland. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: sasp@geus.dk fig. 5. the impressive glaciotectonic thrust sheet exposed in the hanklit coastal cliff. the thrust sheet is 60 m thick and was displaced for a distance of 300 m towards the foreland to the south. mailto:obe@geus.dk geological survey of denmark and greenland bulletin 26, 2012, 65-68 65 new evidence for possible generation of oil off south-western greenland troels laier and hans peter nytoft in 2011, traces of bitumen in the 1160 ma old ilímaussaq intrusion in south greenland have been examined in order to determine their origin. the investigation was prompted by the recent interest in hydrocarbon exploration off western greenland, an interest expressed in the form of four new licences in the region (christiansen 2011). the hydrocarbon potential in the region was realised after reinterpretation of seismic profiles across the labrador sea, and this indicates the presence of a sedimentary basin off south-western greenland (fig. 1; chalmers & pulvertaft 2001). however, the main problem in petroleum exploration off south-western greenland is that no prolific marine source rocks have been demonstrated (christiansen 2011). therefore, any trace of hydrocarbons, however small that may help demonstrate the occurrence of source rocks in the region, deserves careful examination. recently, bitumen biomarkers have been used to question the presumed abiogenic origin of hydrocarbons in crystalline rocks of the ilímaussaq intrusion (laier & nytoft 2012). in this paper, we focus on the origin of the bitumen and compare it with previous finds in central west greenland. the presence of hydrocarbons in the ilímaussaq intrusion has been known since 1970 (petersilie & sørensen 1970) but unlike the discovery of oil seeps in the nuussuaq region in central west greenland twenty years later, which had a positive impact on petroleum exploration (christiansen 2011), the hydrocarbons in the ilímaussaq intrusion were largely ignored in the context of offshore exploration. the reason for this is twofold: (1) hydrocarbons in the ilímaussaq intrusion are much more difficult to recognise than on nuussuaq, and (2) analytical results are confusing with respect to the origin of the hydrocarbons. discrete millimetre-size hydrocarbon accumulations have only been observed twice, and samples of this material were unfortunately not available for analysis in the present investigation. the material, which is a waxy paraffinic hydrocarbon of c28h56, was located in vugs of pegmatite veins and labelled as an evenkitelike mineral by konnerup-madsen et al. (1979). otherwise hydrocarbons in the ilímaussaq intrusion only exist in fluid inclusions, mainly as c1–c5, and as dispersed bitumen invisible to the naked eye. the stable carbon isotopic ratio of methane (δ13c = –7‰) released from the inclusions by crushing (petersilie & sørensen 1970) differed from that of associated methane in most oil and gas reservoirs, which has δ13c values of –30 to –50‰. the ratio was closer to the isotopic ratio of primodial carbon of the earth’s mantle, which has δ13c values around –5‰. the paraffinic hydrocarbons of ‘evenkite’ on the other hand had a δ13c value of –29‰, which is within the expected range for hydrocarbons generated by thermal maturation of organic matter (konnerup-madsen et al. 1988). fig. 1. map of south-western greenland showing the distribution of mesozoic/cenozoic rift basins offshore (green). modified from chalmers & pulvertaft (2001). oil seeps occur on nuussuaq. 100 km greenland 48°w 60°n 66°n 66°n nuussuaq ilímaussaq intrusion marraat © 2012 geus. geological survey of denmark and greenland bulletin 26, 65–68. open access: www.geus.dk/publications/bull 6666 dispersed bitumen in crystalline rocks the distribution of bitumen was examined by microscopy using ultraviolet light, which causes the aromatic constituents of bitumen to fluoresce. unfortunately, a number of the rather common minerals, e.g. sodalite, in different rock types of the ilímaussaq intrusion also fluoresce strongly making it almost impossible to indentify traces of bitumen with certainty in some samples. lujavrite contains only little sodalite and offers the best possibility to study the distribution of bitumen (fig. 2). the bitumen occurs along crystal edges and as trails of tiny inclusions within single crystals of, for example, eudialyte. trails of tiny inclusions are usually taken as evidence of a secondary origin formed in healed fractures. thus hydrocarbons probably migrated through the rock and were trapped in certain minerals. bitumen content and composition it was possible to extract bitumen from naujaite, kakortokite and lujavrite, which are the three major rock types in the ilímaussaq intrusion, by using a 7:1 mixture of dichloromethane and methanol. the bitumen content varied from 110 to 300 mg per kg rock and consisted of paraffins (20%), aromates (20%), nso compounds (compounds with nitrogen, sulphur and oxygen; 50%) and asphaltenes (10%). the nonpolar fraction from the different rock extracts was analysed by gas chromatography with a flame-ionisation detector for total composition and mass spectrometry for biomarker characterisation (fig. 3). all samples gave very similar m/z 217 chromatograms showing a typical marine sterane distribution, although the presence of oleananes in the m/z 191 chromatogram suggests some input from land plants. oleananes are derived from angiosperms, which appeared in the late cretaceous, and hence the presence of oleananes provides a maximum age for the source of the bitumen. since other triterpanes are known to co-elute with oleananes, gas chromatograph mass spectrometry (gc-ms-ms) was also conducted to confirm their presence (fig. 4). the gc-msms analysis not only confirmed the presence of oleananes but also showed the existence of bicadinanes, which is a less common group of biomarkers from land plants from late cretaceous or tertiary. oleananes and bicadinanes were also observed in oil seeps from the nuussuaq region (fig. 4; bojesen-koefoed et al. 1999; nytoft et al. 2002). migration and entrapment of hydrocarbons not only is bitumen much more difficult to recognise in rocks from the ilímaussaq intrusion than in basalts from nuussuaq, but the migration route of hydrocarbons to the ilímaussaq intrusion is also less evident. in nuussuaq the oil seeps are found in tertiary plateau basalts overlying cretaceous and tertiary sediments, some of which are potential a b dc 100 μm fig. 2. photomicrographs of lujavrite (ggu 57033) viewed under visible light (a, b) and in ultraviolet light (c, d) showing hydrocarbons along the edges of eudialyte crystals and as traces of tiny inclusions within the crystals. fig. 3. chromatograms of the aliphatic fraction of extract from ilímaussaq kakortokite m/z 191 (hopanes+tricyclic) and m/z 217 (steranes). fid: flame ionisation detector, gc-ms: gas chromatography mass spectrometry. gc-fid gc-ms m/z 191 pr phy nc15 nc25 ts tm h29 h30 h31 h32 h33 h34 h35 nc30 m/z 217 c29 αα s αα r gc-ms 67 source rocks. in the ilímaussaq intrusion, the bitumen is found in proterozoic crystalline rocks which are much older than the source rock, which is not older than late cretaceous as shown by the presence of oleanes and bicadinanes. in contrast to the nuussuaq region, no potential source rock for hydrocarbons has been reported in south greenland, where proterozoic gneisses and granites, igneous and sedimentary rocks are found. the youngest rocks in south greenland are those of the gardar province (c. 1300–1100 ma), which is dominated by continental sandstones and lavas with numerous dykes and large intrusions (poulsen 1964), one of which is the ilímaussaq peralkaline intrusion. the ilímaussaq intrusion solidified 3–4 km below the surface but is now exposed as a result of erosion. the latest uplift and erosion started c. 35 ma ago according to thermo-chronometric investigations (japsen et al. 2006). however, the latest uplift phase was preceded by subsidence during late cretaceous to eocene. during this subsidence phase the region was probfig. 4. pentacyclic c30 triterpanes in an ilímaussaq bitumen (412 → 191 and 412 → 369). numbered peaks: oleananes (ol + lup) and similar land-plant components (i – iii). peaks h1–h6: hopanoids. peaks t, t1 and r: bicadinanes. 48 h6412 191 100 marraatoil nuussuaq ggu 314654 h1 h2 h3 h4 h5 ol+ lup t i ii iii + ? 43 44 45 4846 47 43 44 45 4846 47 h1 h2 h3 h5 h6 t 412 191 100 ol+ lup h4 h4 h2 t t1 r 412 369 11 43 44 45 4846 47 ol+ lup ol+ lup h1 h2 h3 h4 h5 h6 t 412 191 100 ilímaussaq naujaite ggu 154344 43 44 45 4846 47 ol+ lup h4 t h2 t1 r 412 369 5 43 44 45 46 47 ol+ lup h6412 191 100 412 369 10 43 43 44 44 45 45 48 48 46 46 47 47 h1 h2 h3 h5 t h4 h2 t t1 r h6 ggu 414869a h4 nuussuaq 43 44 45 4846 47 412 369 10 ol+ lup h2 h4 t t1 r i h6 ilímaussaq lujavrite ggu 57033 retention time (min) retention time (min) ol + lup oil seep retention time (min) retention time (min) 6868 ably covered by up to 2 km of marine sediments (chalmers & pulvertaft 2001; japsen et al. 2006). the reburial and uplift history in south greenland was probably similar to that of west greenland, including the nuussuaq region (bonow et al. 2007). thus it is likely that hydrocarbons migrated from marine sediments into the rocks of the ilímaussaq intrusion during the period of reburial (fig. 5). hydrocarbons have only been reported from the ilímaussaq intrusion and not from other rocks in south greenland. there may be two reasons for this: (1) rocks from the ilímaussaq intrusion have been studied in much more detail than other rocks due to the occurrence of rare minerals, some of which have economic potential, and (2) the ilímaussaq intrusion may be more deeply weathered than other rocks in the region because it contains water-soluble minerals (rosehansen & sørensen 2002). from the biomarkers found in the bitumen, there is little doubt that the hydrocarbons were generated by thermal maturation of a marine source rock and migrated into the rocks of the ilímaussaq intrusion at a later stage. it is unlikely that the bitumen formed by condensation of lighter hydrocarbons of abiogenic origin as hypothesised by petersilie & sørensen (1970). hydrocarbon gases in fluid inclusions in rocks from the ilímaussaq intrusion were considered to be abiogenic in origin on the basis of the heavy isotopic value of methane (petersilie & sørensen 1970; konnerup-madsen et al. 1988). however, they may also be of organic origin if the isotopic ratio of the gases was altered by fractionation due to diffusion of gases from the rocks (laier & nytoft 2012). references bojesen-koefoed, j.a., christiansen, f.g., nytoft, h.p. & pedersen, a.k. 1999: oil seepage onshore west greenland: evidence of multiple source rocks and oil mixing. in: fleet, a.j. & boldy, s.a.r. (eds): petroleum geology of northwest europe. proceedings of the 5th conference. petroleum geology conference series 5, 305–314. london: geological society. bonow, j.m., japsen, p., green, p.f., wilson, r.f., chalmers, j.a., klint, k.e., van gool, j.a.m., lidmar-bergström, k. & pedersen, a.k. 2007: a multi-disciplinary study of phanerozoic landscape development in west greenland. geological survey of denmark and greenland bulletin 13, 41–44. chalmers, j.a. & pulvertaft, t.c.r. 2001: development of the continental margins of the labrador sea – a review. in: wilson, r.c.l. et al. (eds): non-volcanic rifting of continental margins: a comparison of evidence from land and sea. geological society special publications (london) 187, 79–107. christiansen, f.g. 2011: greenland petroleum exploration: history, breakthroughs in understanding and future challenges. in: spencer, a.m. et al. (eds): arctic petroleum geology. geological society memoirs (london) 35, 647–661. japsen, p., bonow, j.m., peulvast, j.-p. &wilson, r.w. 2006: uplift, erosion and fault reactivation in southwest greenland. field report summer 2006. danmarks og grønlands geologiske undersøgelse rapport 2006/63, 77 pp. konnerup-madsen, j., larsen, e. & rose-hansen, j. 1979: hydrocarbonrich fluid inclusions in minerals from the alkaline ilímaussaq intrusion, south greenland. bulletin de minéralogie 102, 642–653. konnerup-madsen j., kreulen r. & rose-hansen j. 1988: stable isotopic characteristics of hydrocarbon gases in the alkaline ilímaussaq complex, south greenland. bulletin de minéralogie 111, 567–576. laier, t. & nytoft, h.p. 2012: bitumen biomarkers in the mid-proterozoic ilímaussaq intrusion, southwest greenland – a challenge to the mantle gas theory. marine and petroleum geology 30, 50–65. nytoft, h.p., bojesen-koefoed, j.a., christiansen, f.g. & fowler, m.g. 2002: oleanane or lupane? reappraisal of the presence of oleanane in cretaceous–tertiary oils and sediments. organic geochemistry 33, 1225–1240. poulsen, v. 1964: the sandstones of the precambrian eriksfjord formation in south greenland. rapport grønlands geologiske undersøgelse 2, 16 pp. petersilie, i.a. & sørensen, h. 1970: hydrocarbon gases and bituminous substances in rocks from the ilímaussaq alkaline intrusion, south greenland. lithos 3, 59–76. rose-hansen, j. & sørensen, h. 2002: geology of the lujavrites from the ilímaussaq alkaline complex, south greenland, with information from seven bore holes. meddelelser om grønland, geoscience 40, 58 pp. fig. 5. sketch of the geological evolution of the ilímaussaq area. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: tl@geus.dk including hydrocarbons granite rift deposits lavas intrusion cretaceous palaeogene migrating fluids, a b c d 1150 ma mesozoic palaeogene present geological survey of denmark and greenland bulletin 4, 2003, pp 73-76 one of the first detailed accounts of precambrian supracrustal rocks in central west greenland came from a small group of islands and skerries a few kilometres north-east of aasiaat (fig. 1). in 1948, k. ellitsgaard-rasmussen spent a few days on the islands and published a metamorphic study of their low-grade greenstones and aluminous clastic rocks (ellitsgaard-rasmussen 1954). he observed a striking dissimilarity between these supracrustal rocks and the grey gneisses found in most of the aasiaat region, although the latter were at that time also assumed to be of supracrustal origin. he furthermore noted that the regional significance of the islands should be pursued, and that the island of maniitsoq 4 km west of the small islands might hold a key to their interpretation. more than 50 years were to elapse before the islands were surveyed again in july 2003, during field work for the ikamiut map sheet in the northern nagssugtoqidian orogen (van gool et al. 2002). the collision of two archaean continents during the c. 1850 ma nagssugtoqidian orogeny caused intensive structural and thermal reworking at up to granulite facies grade in most of central west greenland; see connelly et al. (2000) and van gool et al. (2002). the small islands north-east of aasiaat are indeed regionally important, because they document a previously unrecognised low-grade, 73 low-pressure metamorphism during archaean crustal growth: a low-strain zone in the northern nagssugtoqidian orogen, west greenland adam a. garde, mads sylvest christiansen, julie a. hollis, stanislaw mazur and jeroen a.m. van gool fig. 1. geological map of the low-grade, low-strain area north of aasiaat. the well-preserved supracrustal rocks on the island groups around isuamiut and equtiit kangilleq were originally mapped by ellitsgaard-rasmussen (1954). maniitsoq and the small island west of manertooq (a) still preserve intrusive contacts of (?) archaean granodiorite and tonalite into the supracrustal rocks, whereas the intensity of palaeoproterozoic (nagssugtoqidian) strain increases greatly towards the south. the inset map shows the location of the study area in the northern part of the nagssugtoqidian orogen. geological survey of denmark and greenland bulletin 4, 73–76 (2004) © geus, 2004 low-strain domain of presumed archaean age that has largely escaped the nagssugtoqidian orogeny, and as predicted by ellitsgaard-rasmussen (1954) a clue to their significance was found on maniitsoq. the low-grade supracrustal rocks the islands form two different groups (fig. 1). the eastern islands, of which equutit kangilleq is the largest, consist of pale grey, andalusite-staurolite-muscovite-biotite-garnetquartz-bearing metasedimentary rocks interspersed with abundant fine-grained mafic sills. graded bedding on a scale of 5–10 cm is commonly preserved, and up to c. 5 cm large, undeformed andalusite porphyroblasts are found in the aluminous bed tops (fig. 2). randomly orientated, centimetresized staurolite porphyroblasts are also abundant. a 1 m thick aluminous pegmatite with up to 15 cm sized, euhedral andalusite crystals and even larger masses of cordierite were also observed. mafic sills reach thicknesses of a few tens of metres. they are generally fine-grained, massive and dark greenish in colour, and commonly display up to 10 cm thick garnet-bearing reaction rims along their margins. the rocks record a simple deformational history: open to tight folds with amplitudes and wavelengths in the order of 1 km have steeply seor sw-plunging axes and are associated with a steep, centimetre-spaced cleavage that overprints bedding. the coastal exposures of the western islands (of which isuamiut is the largest, fig. 1) consist of dark grey, massive greenstones interspersed with smaller volumes of dark, very fine-grained chlorite schist, mixed horizons of finely layered chert and hematite-dominated, manganiferous banded iron formation (fig. 3) and calcareous layers less than 1 m thick rich in actinolite. most of the greenstones are sill complexes up to c. 1 km thick with common columnar jointing and locally preserved internal intrusive contacts and chilled margins. irregular quartz and calcite veins up to c. 10 cm thick are very common and may have formed by cementation of open joint systems. on the north-west coast of isuamiut a small area of well-preserved pillow lavas was also found (fig. 4), clearly demonstrating that the basic magmatism was contemporaneous with sedimentation, in contrast to the view of ellitsgaard-rasmussen (1954) that the basic magmatism occurred during folding. pillow lava cusps, asymmetric distribution of gas vesicles and graded bedding in adjacent clastic rocks, all point to north-west younging. the general style of deformation, with a simple system of steeply swand seplunging open to close folds, is very similar to that on the eastern islands (fig. 1). steep pencil structures in fold hinges, formed by intersection between cleavage and bedding, demonstrate that the overall deformation was constrictional with subvertical extension. 74 fig. 2. folded aluminous rocks with preserved graded bedding. andalusite porphyroblasts up to 5 cm across (arrows and inset) occur in the bed tops. scale shown by 2.8 cm coin left of centre. south side of small island between isuamiut and equutit kangilleq. fig. 3. vertical layer of banded iron formation and chert, c. 1 m wide, within massive greenstones. northern isuamiut. fig. 4. pillow lava; horizontal surface of steeply dipping unit younging north-west. coin, 2.8 cm across, at cuspate pillow base for scale. northern isuamiut. both the eastern and western islands represent the deposits of a volcanic basin dominated by basic magmatism, with associated chemical sediments now found as chert, banded iron formation and calcareous rocks. the intercalated clastic metasedimentary rocks are very fine-grained and were thus deposited far from continental crustal sediment sources. the origin of the aluminous metasediments on the eastern islands is less clear, although they may also be pelagic sediments. low-grade supracrustal rocks, reminiscent of those described above, occur on hunde ejland and adjacent small islands about 10 km north-west of maniitsoq, but were only briefly surveyed (inset map on fig. 1). layered basic volcanic rocks and sills are intercalated with thin horizons of finegrained chemical and clastic metasedimentary rocks in which chlorite and muscovite are the dominant phyllosilicates; the metamorphic grade appears to have been too low for growth of aluminosilicates. the deformation was sufficiently intense to develop a penetrative schistosity, and bedding-cleavage relationships are only rarely preserved in fold hinges. relationships with the quartzo-feldspathic gneisses in the aasiaat area orthogneisses are absent from the small islands north-east of aasiaat, and their relationships with the low-grade supracrustal association therefore cannot be studied directly. however, a more strongly deformed and higher grade continuation of the supracrustal association probably occurs along strike some 4 km to the west, in the easternmost part of maniitsoq island and on a small island immediately to its east (a on fig. 1). in this area, fine-grained amphibolite considered to be a lateral continuation of the greenstones is intruded by a characteristic unit of k-feldspar megacrystic granodiorite (on maniitsoq), or by grey tonalite (on the small island). the intrusive contacts are weakly deformed but otherwise well preserved (fig. 5). most of maniitsoq is covered by the megacrystic granodiorite, preserved in a weakly deformed state close to its original magmatic appearance; the granodiorite is cut by several sets of flat-lying and inclined pegmatites, and large angles between individual pegmatite phases are still present. a similar, likewise un-migmatised, kfeldspar porphyritic granodiorite was also observed on kronprinsen ejland and may be part of the same pluton. more deformed outcrops of the megacrystic granodiorite have been recognised on several small islands south of maniitsoq, and further south these give way to grey tonalitic orthogneiss (fig. 1). both the granodiorite and orthogneiss become increasingly strongly deformed southwards, and pegmatites are tectonically thinned and lose their angular discordance (fig. 6). southwards, towards aasiaat, the rocks exhibit an intense e–w-trending vertical planar fabric and subhorizontal lineation related to upright, kilometre-scale, tight to isoclinal folds. metamorphism and regional significance the occurrence of andalusite as the stable aluminosilicate phase in staurolite-bearing pelitic rocks is consistent with low pressure metamorphic conditions of ≤ 3 kbar. however, mineral assemblage constraints indicate significant variations in temperature, from chlorite zone greenschist facies up to midupper amphibolite facies. for example, on the western islands chlorite-graphite and garnet-chlorite schists dominate the metasedimentary assemblages; on the eastern islands staurolite-biotite rocks are common, indicating up-temperature crossing of the staurolite isograd (c. 520–550°c). the regional variation in metamorphic grade, coupled with the intrusive relationships observed in the maniitsoq area, strongly suggest metamorphism during emplacement of the granodioritic-tonalitic magmas into the upper crust. the age of emplacement of these plutons is currently unknown; based on regional age data (e.g. connelly et al. 2000) it is presumed that the grey gneisses are late archaean, and u-pb geochronology to confirm this is under way. both previous work in the northern nagssugtoqidian orogen and new observations of intensely deformed palaeo proterozoic basic dykes in the aasiaat region itself, indicate that the intense e–w structural grain in the kangaatsiaq– aasiaat region is due to the nagssugtoqidian continent collision, although almost all of the exposed rocks are of archaean age. contemporaneous deformation also took place north-east of disko bugt, and it has recently been suggested that the nagssugtoqidian orogeny also incorporated the rinkian fold belt in northern west greenland to form a com75 fig. 5. fine-grained amphibolite intruded by weakly deformed tonalite. pen points at angular discordance. west side of island between maniitsoq and manertooq (a on fig. 1). 76 mon, more than 1000 km wide collisional belt extending from west greenland far into eastern canada (thrane et al. 2003). other domains of low palaeoproterozoic strain in the nagssugtoqidian–rinkian orogenic system have previously been described, e.g. from the area north-east of disko bugt (cf. garde & steenfelt 1999) and between kangaatsiaq and attu (piazolo et al. 2004); the former area is generally assumed to consist of a tectonic block that was downthrown along a major extensional shear zone (garde & steenfelt 1999). at aasiaat, however, the strain increase is gradual, taking place over a width of several kilometres. furthermore, the intense subhorizontal lineation in the aasiaat area is perpendicular to the n–s direction of increased strain; it therefore indicates that the main tectonic transport was lateral and did not include a significant vertical component. conclusions greenstones and aluminous metasediments of presumed archaean age crop out on a few small islands north of aasiaat and have been excellently preserved in a low-temperature and low-strain window in the northern part of the nagssugtoqidian orogen in west greenland. these relatively lowgrade rocks may well represent the oldest component of the region, recording a history of metamorphism and deformation during archaean crustal growth. they provide a unique opportunity to study the primary lithological components of the archaean supracrustal belts that are intercalated with the regional grey gneisses. in addition, the state of preservation of the supracrustal rocks provides support for an inhomogeneous nagssugtoqidian orogenic overprint, where blocks with intense thermal and tectonic reworking seem to alternate with blocks of only weak reworking. based on hornblende ar-ar cooling ages, willigers et al. (2002) proposed that the nagssugtoqidian orogeny resulted in uniform heating in most of the orogen (including its northern part), followed by very slow cooling during isostatic uplift. the preliminary observations reported here appear to contradict this, but more work is required to substantiate the new findings. acknowledgements the authors thank christian knudsen, mac persson, sandra piazolo and thomas rasmussen for their contributions to the study of the aasiaat region. references connelly, j.n., van gool, j.a.m. & mengel, f.c. 2000: temporal evolution of a deeply eroded orogen: the nagssugtoqidian orogen, west greenland. canadian journal of earth science 37, 1121–1142. ellitsgaard-rasmussen, k. 1954: on the geology of a metamorphic complex in west greenland. the islands of anarssuit, isuamiut, and eqûtit. bulletin grønlands geologiske undersøgelse 5, 70 pp. garde, a.a. & steenfelt, a. 1999: precambrian geology of nuussuaq and the area north-east of disko bugt. in: kalsbeek, f. (ed.): precambrian geology of the disko bugt region, west greenland. geology of greenland survey bulletin 181, 6–40. piazolo, s., alsop, g.i., nielsen, b.m. & van gool, j.a.m. 2004: the application of gis to unravel patterns of deformation in high grade terrains: a case study of indentor tectonics from west greenland. in: alsop, g.i. & holdsworth, r.e. (eds): flow processes in faults and shear zones. geological society special publication (london) 224, 63–78. thrane, k., connelly, j., garde, a.a. grocott, j. & krawiec, a.w. 2003: linking the palaeoproterozoic rinkian and nagssugtoqidian belts of central west greenland: implications of new u-pb and pb-pb zircon ages. geophysical research abstracts 5(09275). van gool, j.a.m., connelly, j.n., marker, m. & mengel, f.c. 2002: the nagssugtoqidian orogen of west greenland: tectonic evolution and regional correlations from a west greenland perspective. canadian journal of earth science 39, 665–686. willigers, b.j.a., van gool, j.a.m., wijbrans, j.r., krogstad, e.j. & mezger, k. 2002: post-tectonic cooling of the nagssugtoqidian orogen and a comparison of contrasting cooling histories in precambrian and phanerozoic orogens. journal of geology 110, 503–517. fig. 6. intensely deformed granodiorite with several generations of pegmatites rotated into parallelism. the exposure is c. 3 m high. island 2 km east-south-east of manertooq. authors’ addresses a.a.g., m.s.c., j.a.h. & j.a.m.v.g., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: aag@geus.dk s.m., institute of geological sciences, university of wroclaw, pl. maxa borna 9, 50-204 wroclaw, poland. << /ascii85encodepages false /allowtransparency false /autopositionepsfiles true /autorotatepages /all /binding /left /calgrayprofile (dot gain 20%) /calrgbprofile (srgb iec61966-2.1) /calcmykprofile (u.s. web coated \050swop\051 v2) /srgbprofile (srgb iec61966-2.1) /cannotembedfontpolicy /warning /compatibilitylevel 1.4 /compressobjects /tags /compresspages true /convertimagestoindexed true /passthroughjpegimages true /createjdffile false /createjobticket false /defaultrenderingintent /default /detectblends true /colorconversionstrategy /leavecolorunchanged /dothumbnails false /embedallfonts true /embedjoboptions true /dscreportinglevel 0 /syntheticboldness 1.00 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<feff4f7f752890194e9b8a2d5b9a5efa7acb76840020005000440046002065874ef65305542b8f039ad876845f7150cf89e367905ea6ff0c4fbf65bc63d066075217537054c18cea3002005000440046002065874ef653ef4ee54f7f75280020004100630072006f0062006100740020548c002000520065006100640065007200200035002e0030002053ca66f465b07248672c4f86958b555f3002> >> >> setdistillerparams << /hwresolution [2400 2400] /pagesize [595.000 842.000] >> setpagedevice geological survey of denmark and greenland bulletin 38, 2017, 33-36 33 denmark constitutes a low-enthalpy geothermal area. current geothermal production takes place from two sandstone-rich formations: the bunter sandstone and gassum formations. these formations form major potential geothermal reservoirs, but information about the permeability of the potential sandstone reservoirs is difficult to obtain. this may be explained by deposition in a variety of environments under different climatic conditions, and by variable diagenetic overprint (olivarius et al. 2015). thus, the sandstone characteristics and properties are diverse, and in areas where wells are scarce, the assessment of the extent and reservoir properties of sandstone layers is associated with much uncertainty. in order to reduce exploration risk it is therefore essential to develop a robust method for prediction of porosity and permeability prior to drilling. kristensen et al. (2016) presented a five-step method for predicting porosity and permeability averages of prospective geothermal formations in danish onshore areas with low data density. the method is based on the sandstone reservoir characteristics of the gassum formation from geological and petrophysical data acquired in deep wells. the main purpose was to reduce the uncertainties associated with prediction of reservoir properties. the study presented here investigates the efficiency of the five-step method when applied to the bunter sandstone formation. the study area extends from the west coast of denmark to the east coast of southernmost sweden, and from the danish border with germany to just north of the ringkøbing–fyn high (fig. 1). farther north, the bunter sandstone formation grades into the partly contemporary skagerrak formation. eighteen wells provided well-log and core-analysis data from the formation. the five-step method kristensen et al. (2016) developed the five-step method by integrating well log data from the gassum formation with porosity and permeability measurements obtained from conventional core analysis (ccal). the main concept behind the method is that the porosity is related primarily to depth, whereas the permeability depends on a range of parameters including porosity, mineralogy, grain size and sorting. thus, if the burial depth for a particular potential reservoir layer is known, a relatively reliable porosity estimate may be derived (step 1). a permeability estimate may then be obtained from a porosity–permeability relapre-drilling geothermal assessment of porosity and permeability of the bunter sandstone formation, onshore denmark morten leth hjuler and lars kristensen arnum-1 hønning-1 løgumkloster-1–2 varnæs-1 felsted-1 kegnæs-1 tønder-3–5 jelling-1 løve-1 søllested-1 ørslev-1 margretheholm-1 höllviksnäs-1 stenlille-19 rødby-1 50 km danish well swedish well major fault structural high n r i n g k ø b i n g – f y n h i g h n o r t h g e r m a n b a s i n d a n i s h b a s i n f a u l t z o n e 100 km fig. 1. the approximate extent of the bunter sandstone formation in southern denmark and adjacent areas (as delimited by the map frame), showing major structural elements and locations of investigated wells. maps showing e.g. the extent, thickness and data sources of the bunter sandstone formation are available from thewebgis portal at the geological survey of denmark and greenland (http://dybgeotermi.geus.dk/). © 2017 geus. geological survey of denmark and greenland bulletin 38, 33–36. open access: www.geus.dk/publications/bull 3434 tion (steps 2–4), and finally, the uncertainty can be assessed (step 5). therefore the depth of the sandstone layer is required. in areas with no wells, the depth can only be estimated from seismic data, which therefore indirectly control whether the five-step method can be applied. below, the main components of the five steps are defined and explained; for a detailed description of the concept, refer to kristensen et al. (2016). step 1: porosity–depth model. in order to predict the porosity of a formation, a regional porosity–depth relation is established. log-derived, effective porosity data subjected to cut-off by shale content (<30%) and porosity (>15%) are averaged for each well and plotted against estimated maximum burial depth. the resulting relation, the porosity– depth model, is expressed as ϕlog = a + b × d (equation 1), where ϕlog is the log-derived porosity, d is the burial depth and a and b are constants. step 2: initial permeability model. an initial, basin-wide porosity–permeability relation based on core-analysis data is established in order to predict permeability. subsequent to porosity cut-off (15%), this relation, i.e. the initial permeability model, is expressed as kini = a × ϕcore b (equation 2), where kini is the core-gas permeability, ϕcore is the core porosity and a and b are constants. step 3: general permeability model. this model uses logderived, averaged permeabilities with the purpose of incorporating a dataset encompassing the entire formation and not only covering parts of the formation as is mostly the case with the core-based dataset used in step 2. equation 2 is used to calculate permeability estimates from the log-derived porosity. shale and porosity cut-offs of 30% and 15%, respectively, are applied. a cross-plot between log-derived and averaged porosity and permeability data points forms the basis for definition of a general permeability model expressed as kg = a × ϕlog b (equation 3), where kg is the log-derived permeability, ϕlog is the log-derived porosity and a and b are constants. step 4: local permeability model. the general permeability model is adapted to local conditions by multiplying equation 3 with a constant, c. thus, the general permeability model constitutes a template for the local permeability model, which may be expressed as kl = c × kg (equation 4), where kl is the local permeability. step 5: permeability-uncertainty range. a local area featuring several wells with a sufficient amount of appropriate log data is selected and a local permeability model is established. the extensive source data ensure a statistically well-restrained model and a consistent uncertainty range, which is assumed to be applicable to local models within the same formation in other parts of the danish area. below, the results of applying the five-step method to the bunter sandstone formation in denmark are described. application of the five-step method fourteen wells (fig. 1) provided data for the generation of a porosity–depth model for the bunter sandstone formation (fig. 2); data from the løgumkloster-1, 2 and 2a wells were excluded because of their location in a fault zone (fig. 1) where tectonic activity may have altered the sandstone properties. the porosity–depth model (step 1). unfortunately, the scattered data distribution prevented derivation of a mathematically defined trend line, and thus a reliable porosporosity % rødby-1 felsted-1 kegnæs-1 kværs-1 margretheholm-1 stenlille-19 varnæs-1 løve-1 ørslev-1 jelling-1 søllested-1 løgumkloster-2a løgumkloster-2 løgumkloster-1 tønder-3 tønder-4 tønder-5 margretheholm-1 stenlille-19 løgumkloster-1 tønder-4 tønder-5 average net porosity porosity–depth trend gluyas & cade (1997) 1500 2000 2500 3000 3500 0 10 20 30 es tim at ed   m ax im al   bu ria l  de pt h  (m ) fig. 2. porosity–depth model (step 1) for the bunter sandstone formation. the shown model (black line) is an estimate by the authors; the løgumkloster wells were excluded from the model (see text) due to their location in a fault zone. a mechanical compaction curve (gluyas & cade 1997) is shown for comparison. modified from kristensen et al. (2016). 0 10 20 30 40 50 1 10 100 1000 10000 g as  p er m ea bi lit y  (m d ) porosity % initial permeability model: κini = 48084 × φcore 3.4571 arnum-1 hønning-1 höllviksnäs-1 rødby-1 stenlille-19 tønder-3 tønder-4 tønder-5 porosity cut-off fig. 3. initial permeability model (step 2) for the bunter sandstone formation. 35 ity–depth model could not be generated. the shown trend line is the authors’ best estimate based on their general experience and knowledge of petrographic characteristics and diagenetic overprints. the generation of a permeability–depth model was attempted as an alternative to the porosity–depth model, but without success as the scattered data distribution prevented establishment of a well-defined trend line. the initial permeability model (step 2). eight of the wells shown in fig. 1 provided core-analysis data for the initial permeability model (fig. 3). in order to generate a reliable porosity–permeability relation based on comparable measurements, only permeabilities obtained from cleaned, unflawed plugs of productive lithologies, measured under similar testing conditions were used. data from uncleaned plugs and data points representing low permeabilities (<1 md) and unproductive lithologies (claystone/shale) were removed from the dataset. subsequently, shale and porosity cut-offs of 30% and 15%, respectively, were applied. the initial permeability model is based on 360 data points and is expressed by equation 5: kini = 48084 × ϕcore 3.4571, where kini is in md and ϕcore is in per cent. the general permeability model (step 3). the log-derived permeability was calculated for 16 wells (fig. 1) using equation 5 with the log-derived porosity as input. løgumkloster-2a was discarded due to porosity cutoff. subsequently, the log-derived permeability was averaged and a cross-plot between averaged log-derived porosity and permeability was generated (fig. 4). the general permeability model is expressed by equation 6: kg = 0.0026 × ϕlog 3.7843 (fig. 4). the introduction of the general permeability model imposes a slightly higher specific permeability estimate than that calculated from the initial permeability model, i.e., kg > kini (fig. 4). the local permeability model (step 4). local permeability models were generated for the tønder and stenlille areas in order to demonstrate the variability of sandstone reservoirs within the bunter sandstone formation (fig. 5). for the tønder area, the local permeability model is expressed as ktønder = 1.7 × (0.0026 × ϕ3.7843) (equation 7) and for the stenlille area as kstenlille = 0.8 × (0.0026 × ϕ3.7843) (equation 8). permeability-uncertainty range (step 5). only the tønder area provides the data density needed for establishing an uncertainty range for a local permeability model. however, the number of data-supplying wells is limited to 3 (tønder-3–5), and the statistical basis from the available bunter sandstone formation data is insufficient to assess the uncertainty range. results and discussion porosities of the bunter sandstone formation are lower than derived from the mechanical compaction curve for uncemented sandstones (step 1, fig. 2), indicating that most of these sandstones contain clay and diagenetic cement (kristensen et al. 2016). at greater depths, the sandstones preserve relatively high porosity due to microporosity within detrital clays and diagenetic iron oxide/ hydroxide coatings that seemingly retard quartz cementation (olivarius et al. 2015). although a well-defined porosity–depth model (step 1) could not be created for the bunter sandstone formation, the range of log-derived average porosities from 17% to 25% within the 1500–3100 m depth interval indicates favourable porosity conditions irrespective of maximum burial depth (fig. 2). a corresponding averaged permeability range was calculated using the general permeability model (step 3, equation 6) with 17% and 25% as porosity inputs. within these bounds the permeability ranges from 132–557 md (fig. 6). from this observation, the geothermal prospectivity of the bunter sandstone formation is significant, provided sufficient reservoir thickness and temperature are present. 50 10 15 20 25 30 1 10 100 1000 10000 g as  p er m ea bi lit y  (m d ) porosity % log data: bunter sandstone fm general permeability model porosity cut-off initial permeability model: general permeability model:     κini = 48084 × φcore 3.4571 κg = 0.0026 × φlog 3.7843 core-analysis data: initial permeability data fig. 4. initial and general permeability models (steps 2, 3) for the bunter sandstone formation after averaging the log-derived permeability and porosity. 50 10 15 20 25 30 1 10 100 1000 10000 g as  p er m ea bi lit y  (m d ) core-analysis data: stenlille 19 (fm average) stenlille 19 tønder 3 (fm average) tønder 3 tønder 4 (fm average) tønder 4 tønder 5 (fm average) tønder 5 log data: general model  tønder model   stenlille model  porosity cut-off local permeability models: κtønder = 1.7 × 0.0026 × φ3.7843 κstenlille = 0.8 × 0.0026 × φ3.7843 porosity % fig. 5. local permeability models (step 4) for the bunter sandstone formation in the tønder and stenlille areas. 3636 the vast majority of data behind the initial permeability model (kini, step 2) belong to the tønder-3–5 wells. therefore, kini represents the tønder area rather than the entire study area. this data distribution issue was resolved by applying the averaging technique, whereby each well is represented by a single data point in the porosity–permeability plot. thus, by using averaged data, each well becomes equally influential in the general permeability model (kg, step 3), see fig. 4. further benefits of averaging the data are a dramatic reduction of the scatter of core-analysis measurements (fig. 3) and a significant narrowing of the uncertainty band associated with average permeability estimates. in step 4, the local permeability models generated for the tønder and stenlille areas, ktønder and kstenlille, demonstrate the importance of including local geological data when establishing local permeability models. at any porosity, ktønder is 1.7 times higher than kg and more than twice that of kstenlille. in contrast, kstenlille is 0.8 times lower than kg. these permeability variations may be explained by variations in deposition environment and maximum burial depth. at tønder, aeolian deposition favoured the generation of well-sorted, clay-free sandstones which, subsequent to maximum burial of c. 2000 m, resulted in high average porosities and permeabilities (c. 22% and 300 md, respectively); the presence of nitrogen gas may play a porositypreserving role. at stenlille, the sandstones were deposited in alluvial fan and braided river systems (olivarius & nielsen 2016) and maximum burial occurred at c. 3000 m, i.e. 1000 m deeper than at tønder (fig. 2), possibly causing the lower porosity (c. 19%) and permeability (c. 180 md) due to poorer sorting and a higher degree of mechanical compaction and diagenesis. it was not possible to establish a permeability uncertainty band for the bunter sandstone formation (step 5). however, assuming that the bunter sandstone and gassum formations respond similarly to the geological factors controlling permeability variations, it is suggested that the uncertainty range derived for the gassum formation by kristensen et al. (2016) is also applicable to the bunter sandstone formation. therefore, the uncertainty range may be expressed by multipliers of 2 and 0.5. conclusions it was not possible to generate a reliable porosity–depth model (step 1) for the bunter sandstone formation based on data and modelling from the gassum formation. however, using averaged porosities and permeabilities from all available wells (step 3), porosities from 17–25% and permeabilities from 132–557 md can be modelled in the depth range 1500–3100 m. local permeability models (step 4) for the bunter sandstone formation in the tønder and stenlille areas differ significantly from the general permeability model (step 3), which emphasises the importance of using local geological data for calibration of the general permeability model. a permeability uncertainty band (step 5) could not be established for the bunter sandstone formation due to insufficient data; however, the uncertainty range developed for the gassum formation in kristensen et al. (2016) may be applied. references gluyas, j. & cade, c.a. 1997: prediction of porosity in compacted sands. in: kupecz, j.a., gluyas j. & block, s. (eds): reservoir quality prediction in sandstones and carbonates. aapg memoir 69, 19–28. kristensen, l., hjuler, m.l, frykman, p., olivarius, m., weibel, r., nielsen, l.h. & mathiesen, a. 2016: pre-drilling assessments of average porosity and permeability in the geothermal reservoirs of the danish area. geothermal energy 4:6, 27 pp. olivarius, m. & nielsen, l.h. 2016: triassic paleogeography of the greater eastern norwegian-danish basin: constraints from provenance analysis of the skagerrak formation. marine and petroleum geology 69, 168–182. olivarius, m., weibel, r., hjuler, m.l., kristensen, l., mathiesen, a., nielsen, l.h. & kjøller, c. 2015: diagenetic effects on porosity– permeability relationships in red beds of the lower triassic bunter sandstone formation in the north german basin. sedimentary geology 321, 139–153. 50 10 15 20 25 30 1 10 100 1000 10000 g as  p er m ea bi lit y  (m d ) porosity % averaged log data: bunter sandstone fm general permeability model porosity cut-off 132 md 557 md  general permeability model:    κg = 0.0026 × φlog 3.7843 fig. 6. porosity and permeability range for the bunter sandstone formation in the investigated wells in denmark. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: mlh@geus.dk geologial survey of denmark and greenland bulletin 4, 2003, p 65-68 65 in 2003, the geological survey of denmark and greenland (geus) completed a four-year project aimed at assessing the mineral potential of the precambrian region of west greenland between latitudes 66° and 70°15´n. the project was part of a contract between geus and the ministry of the environment, and involved compilation of existing geoscientific data, new geological mapping, field examinations of known and potential mineral occurrences, new chemical and isotope analyses, and data interpretation. the data compilation, available on a dvd (schjøth et al. 2004), comprises regional, systematically acquired data sets presented in a geographic information system environment. aeromagnetic, aeroradiometric, stream sediment and rock geochemical and gravity data, a digital elevation model and a satellite image are included, plus descriptions of 60 mineral occurrences. evaluation of the mineral potential is based on interpretations of the compiled information as well as on earlier investigations by the survey, the university of copenhagen and commercial companies (see e.g. stendal & schønwandt 2003; stendal et al. 2004). from an economic point of view, the potential for gold and diamonds is the most interesting in the investigated area. this paper summarises the evaluation of the gold potential; results of diamond-related investigations are reported separately (jensen & secher 2004, this volume). geological setting while the ataa and nassuttooq areas (fig. 1) were comparatively well known before the onset of the project, only limited information was available on large parts of the intervening tract. important aspects of the present project have therefore included establishment of the geological relationships between the northern and southern areas, as well as the ages and plate-tectonic settings of lithological units and their mineralisation. an overview of the geology of the area is presented in fig. 1. geological descriptions and further references may be found in garde & steenfelt (1999) and van gool et al. (2001). granitoid orthogneisses of archaean age with subordinate supracrustal sequences underlie almost the entire area. radiometric age determinations and structural studies have gold in central west greenland – known and prospective occurrences agnete steenfelt, henrik stendal, bo møller nielsen and thorkild m. rasmussen fig. 1. map of central west greenland showing supracrustal rocks and known gold occurrences. isotopic data suggest that some supracrustal sequences comprise both archaean and palaeoproterozoic rocks (question marks). cno, central nagssugtoqidian orogen; ape, arveprinsen ejland. modified from van gool et al. (2002). geological survey of denmark and greenland bulletin 4, 65–68 (2004) © geus, 2004 documented that some of the supracrustal sequences are archaean while others are palaeoproterozoic in age. most of the archaean crust in the study region was formed around 2.8 ga, probably by accretion of tonalitic magmatic complexes formed at constructive continental margins. however, remnants of older continents have also been recognised, such as enclaves of mid-archaean (c. 3 ga) rock complexes and a small occurrence of early archaean orthogneiss (3.6 ga) in the south-eastern corner of the region. the metavolcanic sequences in the ataa area were probably formed in a volcanic arc setting, while supracrustal sequences elsewhere are dominated by sediments and are likely to represent continental rift or margin settings. late archaean (c. 2.75 ga) crustal heating with formation of pegmatites and granite veins has affected the entire region. the archaean basement together with palaeoproterozoic cover rocks has been variably affected by the palaeoproterozoic nagssugtoqidian orogeny that lasted from c. 2.0 to 1.75 ga. the orogeny involved initial continental rifting accompanied by intrusion of a dolerite dyke swarm, a subduction phase with extrusion and intrusion of quartz dioritic and monzodioritic magmas, and a final continental collision phase around 1.84 ga resulting in folding, thrusting, shearing, metamorphism and local melting. at the same time as post-kinematic pegmatites and granites were emplaced at c. 1.8 ga north of the central nagssugtoqidian domain, ultramafic lamprophyres were intruded in the ataa area. much later, the southern continent of the nagssugtoqidian orogen was the focus of recurrent alkaline and carbonatitic magmatism (larsen & rex 1992) resulting in numerous lamprophyre and carbonatitic dykes (see jensen & secher 2004, this volume) in addition to the major sarfartoq carbonatite complex (c. 0.6 ga). known gold occurrences eight mineral occurrences with prospective gold concentrations (i.e. above one gramme per ton, 1 ppm) have been identified during commercial and government exploration in the study region, although none have so far proved economic in size and grade (fig. 1; stendal 1998; stendal et al. 2002; stendal & schønwandt 2003). the known auriferous mineral occurrences are all hosted in archaean rocks. two are stratabound and were probably formed during deposition of the host sediments and lavas, while the others are located in cross-cutting veins, alteration zones, breccias and shear zones. itilli. the oldest known metavolcanic rocks in the region are the amphibolites at itilli, which are intruded by a 3 ga old quartz diorite (thrane & connelly 2002). the setting is possibly a volcanic arc. disseminated sulphides within the amphibolites are closely related to short, variably oriented shear zones. locally these shear zones host quartz lenses (5–10 cm wide and 1 m long) parallel to the fabric of the shear zone, that are often surrounded by a halo of hydrothermal alteration. veinlets and disseminated sulphides including chalcopyrite occur both in the quartz lenses and the sheared amphibolite and contain 0.7 ppm au. other shear zones along a thrust contact between amphibolites and metasediments, up to 25 cm wide, contain quartz lenses and nickelarsenides (e.g. gersdorffite). analyses of chip samples from these shear zones have yielded values of up to 0.3% cu, 1.3 ppm au and 1.5% ni. saqqaq. the supracrustals comprise felsic metasedimentary rocks with subordinate mafic and ultramafic metavolcanic units. the sequence may represent deposition in a continental rift or at an active continental margin. the age is not known, but believed to be archaean. gold occurs in a 1 to 2 m thick garnet-quartz horizon situated at the boundary between ultramafic lavas and mica schist. the horizon has been interpreted as a syngenetic exhalative chert (garde et al. 1999), or alternatively a silicified shear zone (nunaminerals 2000). the auriferous metachert layer contains a few per cent of disseminated sulphides. gold values are in the range of 1–16 ppm over 1–2 m, and the auriferous bed can be followed for at least 4 km. in addition to gold, the mineralised layer also has high concentrations of as (average 404 ppm), ni (average 652 ppm) and cr (average 1403 ppm). eqi. the metamorphosed mafic and felsic volcanic rocks are considered part of a 2.8 ga old volcanic island arc. three types of gold mineralisation are found. syngenetic gold occurs in up to 20 cm thick lenses of semi-massive pyrite situated in a 50–200 m wide zone between rhyolitic lava and sericite-rich sediment. composite grab samples of massive pyrite have yielded up to 0.2% cu and 1 ppm au. hydrothermal activity associated with the volcanism at eqi resulted in pervasive carbonate alteration along n–s-trending zones. the carbonatised rocks consist of ankerite, chlorite, green fuchsitic mica and disseminated pyrite. chip samples of carbonate alteration have given up to 2.3 ppm au over 2.5 m, while grab samples of quartz-veined rocks have yielded between 5 ppb (microgramme per ton) and 60 ppm au. the third kind of gold mineralisation is hosted by a 10 m wide and 100 m long breccia zone, is clearly epigenetic and is palaeoproterozoic in age. the breccia is situated immediately west of a major n–s-trending thrust separating the greenschist facies volcanic rocks to the east from amphibolite facies sedimentary rocks to the west. the breccia includes pyrrhotite, pyrite and chalcopyrite, and minor amounts of sphalerite and arsenopyrite. the main zone yielded up to 1.7 ppm au over 1 m, and the best section of 12 short drill holes assayed 1.3% cu and 12 ppm au over 3.2 m. 66 arveprinsen ejland. the supracrustal sequence (fig. 1, ape) probably forms part of the same volcanic island arc represented at eqi. it comprises mafic metavolcanic lavas and a mafic sill complex. disseminated pyrrhotite and pyrite are common and semi-massive to massive lenses of sulphide also occur. quartz veins with up to 2.6 ppm au occur in minor pyrite-rich shear zones. samples of hydrothermally altered sulphide within a fault zone have assayed between 8 and 25 ppm au. pb-pb isotope data indicate that the hydrothermal alteration is palaeoproterozoic in age. itilliarsuk. the metasediment-dominated supracrustal sequence along the south coast of nuussuaq hosts large, rusty-weathering iron-formations in the form of both magnetite-rich bands and semi-massive sulphide. mica schist with disseminated pyrrhotite and pyrite attains a thickness of about 150 m. nunaminerals (2000) have identified several sites with epigenetic gold mineralisation in sulphide-rich schists, quartz veins and shear zones. the best target is a shear zone that hosts a quartz-sericite rock yielding 9 ppm of au over 1.7 m. the mineralised structure can be traced 500 m along strike. attu. this gold occurrence is the only place outside the ataa area where gold concentrations above 1 ppm have so far been recorded. a prominent mylonite zone comprising pegmatite veining, silicification, magnetite and sulphide mineralisation cuts through granulite facies orthogneiss. the zone is 100–330 m wide, strikes 75° and dips 60–70°w. samples of silicified mylonite assayed 2 to 8 ppm au. favourable areas for further gold prospecting the distribution and ages of known gold occurrences within the study region demonstrate that gold-bearing mineralising systems have been active at various times throughout the geological evolution of the study region. syngenetic gold mineralisation is evident in archaean times where the favourable environments appear to have been rift or active continental margin and volcanic island arcs. renewed heating and associated hydrothermal circulation in these environments could remobilise and deposit gold where fluids moving along shear zones and faults pass through basic, reducing rocks at the appropriate temperature. several events are known, which may have triggered epigenetic gold mineralisation: c. 2.75 ga granite intrusions, palaeoproterozoic (c. 1.9 ga?) intrusion of sills in the ataa area, c. 1.84 ga deformation and heating resulting from continent collision, 1.8 to 1.7 ga granite and pegmatite veining, c. 1.75 ga lamprophyre intrusion, and albitisation in the ataa area, possibly younger than 1.7 ga. figure 2 shows favourable areas for gold mineralisation based on the following criteria: (1) presence of stream sed67 fig. 2. gold anomalies and gold prospective areas in central west greenland. stream sediment collected over the entire region at a spacing of 4 to 6 km by geus (steenfelt 2001). heavy mineral concentrates of stream sediment (hmc) collected in selected areas by nunaoil a/s (sieborg 1992). hmc and < 0.1 mm grain size fraction of stream sediment types analysed by the instrumental neutron activation method. the distribution of high concentrations of as, sb and rb is illustrated by combined contoured grids. gridding was carried out with a cell size of 5 km using a kriging method. the triangle shows range of colours where grids overlap. nssz, nordre strømfjord shear zone. 68 ments with high gold contents in the fine fraction or in heavy mineral concentrates; (2) areas where high arsenic (as), antimony (sb) and rubidium (rb) in stream sediment data coincide. as and sb are frequently mobilised and precipitated together with au in nature, and high rb reflects the distribution of granitic magmatism; (3) areas in the vicinity of major faulting, shearing and thrusting; (4) presence of supracrustal rocks (see fig. 1), particularly basic and reducing lithologies. apart from the ataa area where conditions are very favourable (fig. 1), the area around naternaq seems the most prospective. gold values of c. 2 ppm have been reported, but such concentrations have not been reproduced in samples collected during the present project. the same is the case regarding the cluster of anomalies near kangersuneq. later prospecting and sampling by nunaminerals and geus have been unable to explain or reproduce the anomalies. other environments of interest for further prospecting are the nordre strømfjord shear zone (nssz), and the vicinity of major thrust zones in the southern study region (fig. 2). references garde, a.a. & steenfelt, a. 1999: precambrian geology of nuussuaq and the area north-east of disko bugt, west greenland. geology of greenland survey bulletin 181, 7–40. garde, a.a., thomassen, b., tukiainen, t. & steenfelt, a. 1999: a goldbearing volcanogenic-exhalative horizon in the archaean(?) saqqaq supracrustal rocks, nuussuaq, west greenland. geology of greenland survey bulletin 181, 119–128. jensen, s.m. & secher, k. 2004: investigating the diamond potential of southern west greenland. geological survey of denmark and greenland bulletin 4, 69–72 (this volume). larsen, l.m. & rex, d.c. 1992: a review of the 2500 ma span of alkalineultramafic, potassic and carbonatitic magmatism in west greenland. lithos 28, 367–402. nunaminerals 2000: an overview of the company and its prospects, 8 pp. unpublished report, nunaminerals a/s, nuuk. schjøth, f., & steenfelt, a., (eds) 2004: mineral resources of the precambrian shield of central west greenland (66° to 70°15´n). part 1. compilation of geoscience data. danmarks og grønlands geologiske undersøgelse rapport 2004/16, 45 pp. sieborg b. 1992: geochemical exploration in west greenland. july–august 1991, vol. 1, 36 pp., 6 app. vol. 2, 28 plates. unpublished report, nunaoil a/s, copenhagen, denmark (in archives of geological survey of denmark and greenland, geus report file 21080). steenfelt, a. 2001: geochemical atlas of greenland – west and south greenland. danmarks og grønlands geologiske undersøgelse rapport 2001/46, 39 pp. + 1 cd-rom. stendal, h. 1998: contrasting pb isotopes of archaean and palaeo-proterozoic sulphide mineralisation, disko bugt, central west greenland. mineralium deposita 33, 255–265. stendal, h. & schønwandt, h.k. 2003: precambrian supracrustal rocks and mineral occurrences, northeast disko bugt. danmarks og grønlands geologiske undersøgelse rapport 2003/24, 57 pp. stendal, h., blomsterberg, j., jensen, s.m., lind, m., madsen, h.b., nielsen, b.m., thorning, l. & østergaard, c. 2002: the mineral resource potential of the nordre strømfjord – qasigiannguit region, southern and central west greenland. geology of greenland survey bulletin 191, 39–47. stendal, h., nielsen, b.m., secher, k. & steenfelt, a. 2004: mineral resources of the precambrian shield of central west greenland (66° to 70°15´n). part 2. mineral resources. danmarks og grønlands geologiske undersøgelse rapport 2004/20, 212 pp. thrane, k. & connelly, j. 2002: linking the nagssugtoqidian orogen and the rinkian belt: preliminary ages from the disko bugt region. in: nielsen, b.m. & thrane, k. (eds): workshop on nagssugtoqidian and rinkian geology, west greenland. danmarks og grønlands geologiske undersøgelse rapport 2002/9, 46–48. van gool, j.a.m., connelly, j.n., marker, m. & mengel, f. 2001: the nagssugtoqidian orogen of west greenland: tectonic evolution and regional correlations from a west greenland perspective. canadian journal of earth sciences 39, 665–686. van gool, j.a.m., alsop, i., árting, u.e., garde, a.a., knudsen, c., krawiec, a.w., mazur, s., nygaard, j., piazolo, s., thomas, c.w. & thrane, k. 2002: precambrian geology of the northern nagssug toqidian orogen, west greenland: mapping in the kangaatsiaq area. geology of greenland survey bulletin 191, 13–23. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ast@geus.dk << /ascii85encodepages false /allowtransparency false /autopositionepsfiles true /autorotatepages /all /binding /left /calgrayprofile (dot gain 20%) /calrgbprofile (srgb 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<feff4f7f752890194e9b8a2d5b9a5efa7acb76840020005000440046002065874ef65305542b8f039ad876845f7150cf89e367905ea6ff0c4fbf65bc63d066075217537054c18cea3002005000440046002065874ef653ef4ee54f7f75280020004100630072006f0062006100740020548c002000520065006100640065007200200035002e0030002053ca66f465b07248672c4f86958b555f3002> >> >> setdistillerparams << /hwresolution [2400 2400] /pagesize [595.000 842.000] >> setpagedevice geological survey of denmark and greenland bulletin 26, 2012, 1-8 1 geological survey of denmark and greenland bulletin 26 • 2012 review of survey activities 2011 edited by ole bennike, adam a. garde and w. stuart watt geological survey of denmark and greenland ministry of climate, energy and building 22 geological survey of denmark and greenland bulletin 26 keywords geological survey of denmark and greenland, survey organisations, current research, denmark, greenland. cover photographs from left to right 1. retrieving a sediment core in sermilik fjord, south-east greenland. photograph: robert s. fausto. 2. field work in vietnam. photograph: flemming larsen. 3. this danish beech forest on fyn was farm land during the iron age. photograph: ole bennike. 4. small-scale miners in nigeria. worldwide, c. 100 million people depend on small-scale mining. photograph: john tychsen. frontispiece: facing page in the summer of 2011, geus carried out extensive reconnaissance work in south-east greenland. the aim is to assess the mineral potential of the region, which is one of the least known regions of greenland. the geologist is panning stream sediments in order to separate heavy minerals, possibly including gold. photograph: jakob lautrup. chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, department of geoscience, aarhus university; minik rosing, geological museum, university of copenhagen; finn surlyk, department of geography and geology, university of copenhagen scientific editors: ole bennike, adam a. garde and w. stuart watt editorial secretaries: jane holst and esben w. glendal referees: (numbers refer to first page of reviewed article): johanna anjar, se (29); anonymous (65), niels balling, dk (41); stefan bernstein, dk (57); albertas bitinas, lt (29); lars ole boldreel, dk (81); henrik breuning, dk (85); gregers dam, dk (61); margaret dolan, no (25); ida fabricius, dk (9); tomas feseker, de (69); kerstin geitner, dk (25); lawrence gill, dk (65); sam holloway, uk (45); michael houmark-nielsen, dk (17); shfaqat abbas khan, dk (41); poul-henrik larsen, dk (61); james lawrence, uk (33); clive mitchell, uk (85); farrokh nadim, no (33); thomas pape, de (69); heikki papunen, fi (53); asger ken pedersen, dk (57); gunver krarup pedersen, dk (9); stefan piasecki, dk (13); hans plaat, nl (45); manuel pubellier, fr (81); jan audun rasmussen, dk (13); anders schomacker, no (17); inga sørensen, dk (37); jette sørensen, dk (37); mathilde b. sørensen, no (77); henrik stendal, gl (53); kristian syberg, dk (49); matti tarvainen, fi (77); henry vallius, fi (21); karen vilholth, lk (49); michiel van den broecke, nl (73); roy h. wilkens, usa (21); jacob clement yde, no (73). illustrations: stefan sølberg, jette halskov, willy weng, frants v. platen-hallermund and benny m. schark layout and graphic production: kristian rasmussen printers: rosendahls-schultz grafisk a/s, albertslund, denmark manuscripts received: 6 january 2012 – 8 may 2012 final versions approved: 13 february 2012 – 23 may 2012 printed: 10 july 2012 issn 1604-8156 isbn 978-87-7871-339-1 citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of the name of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 26, 88 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk. see also www.geus.dk/publications/bull © de nationale geologiske undersøgelser for danmark og grønland (geus) 2012 for the full text of the geus copyright clause, please refer to www.geus.dk/publications/bull 3 44 namibia botswana south africa zambia ghana tanzania mozambique greenland kenya cameroun ethiopia nigeria bolivia brazil mexico denmark 7 review of survey activities 2011 f.g. christiansen 9 nano-quartz in north sea danian chalk h. lindgreen and f. jakobsen 13 geology of the femern bælt area between denmark and germany e. sheldon, p. gravesen and h. nøhr-hansen 17 rock-cored drumlins on bornholm, denmark p.r. jakobsen 21 methane distribution in holocene marine sediments in the bornholm basin, southern scandinavia j.b. jensen and r. endler 25 natura 2000 habitat mapping in kattegat, denmark: an example from læsø trindel z.k. al-hamdani and l.g. addington 29 early holocene sea-level changes in øresund, southern scandinavia o. bennike, m.s. andreasen, j.b. jensen, m. moros and n. noe-nygaard 33 cliff collapse at stevns klint, south-east denmark s.a.s. pedersen and t. damholt 37 shallow geothermal energy in denmark t. vangkilde-pedersen, c. ditlefsen and a.l. højberg 41 eff orts to include geological and geodetic observations in the assessment of earthquake activity in denmark s. gregersen and p.h. voss 45 results of monitoring groundwater above the natural gas underground storage at stenlille, denmark t. laier 49 groundwater protection in denmark and the role of water supply companies j.d. petersen and l.f. jørgensen 5 india indonesia vietnam the philippines thailand geus working areas 2011. orange areas are covered in this volume. 53 anorthosites in greenland: a possible raw material for aluminium? c. knudsen, j. wanvik and h. svahnberg 57 from 3d mapping to 3d modelling: a case study from the skaergaard intrusion, southern east greenland k. svennevig and p. guarnieri 61 geological assessment of the east greenland margin m.b.w. fyhn, t.m. rasmussen, t. dahl-jensen, w.l. weng, j.a. bojesen-koefoed and t. nielsen 65 new evidence for possible generation of oil off south-western greenland t. laier and h.p. nytoft 69 methane and possible gas hydrates in the disko bugt region, central west greenland n. mikkelsen, t. laier, t. nielsen, a. kuijpers and n. nørgaard-pedersen 73 ablation observations for 2008–2011 from the programme for monitoring of the greenland ice sheet (promice) r.s. fausto, d. van as and the promice project team 77 testing of an automatic earthquake detection method on data from station nord, greenland n. karamzadeh, p.h. voss and g.d. javan 81 th e cenozoic song hong and beibuwan basins, vietnam m.b.w. fyhn, h.i. petersen, l.h. nielsen, t.c. giang, l.h. nga, n.t.m. hong, n.d. nguyen and i. abatzis 85 rock phosphate and lime for small-scale farming in tanzania, east africa p. kalvig, n. fold, j.b. jønsson and e.e. mshiu 66 7 review of survey activities 2011 flemming g. christiansen deputy director 2011 was a good year for the geological survey of denmark and greenland (geus) with fruitful discussions of strategies, handling of challenges posed to our society that has strong needs for growth and investments, and consequently for developing new projects for the future within the many fields where geus is involved. geus has been through a long – but very constructive – process of putting together a new strategy that reflects the changes in society and new demands from many different stakeholders. this new strategy, ‘geology for society – knowledge for growth and welfare’, covers the period up to 2020 and is based on an analysis of key driving forces such as transition from fossil fuel to sustainable energy, adaptation to changing climate conditions, need for water resources locally and globally, and international competition for critical minerals. the strategy includes a number of specific strategic topics. as these topics have been discussed between scientists and management over a long period and in some cases are related to new and on-going survey projects, these topics are reflected in many of the papers in this year’s issue of review of survey activities. the reader will find contributions covering several of our new strategic topics such as: ‘past and future climate’, ‘geology across land and sea’, ‘geology in the public arena’, ‘water resources under pressure’, ‘oil and gas supply and transition to green energy’ and ‘geothermal energy and heat storage’, as well as papers on other subjects from geus’ very broad project portfolio. some of them describe specific, short projects, and others focus on results from some of our long-term monitoring projects on for example groundwater, gas storage, ice and earthquakes. geus’ activities and research in denmark cover many different subjects such as fossil fuel and renewable energy, water, climate history and monitoring and adaptation to climate changes, nature protection, input to large infrastructure projects and data management. chalk is the main oil reservoir in the danish part of the north sea and understanding its mineralogy and diagenesis has crucial implications for production parameters. one paper proposes a new model for the formation of flint from studies of nano-quartz. as part of planning the fixed link across femern bælt between denmark and germany, geus has been involved in geological and geotechnical investigations. results of a multidisciplinary biostratigraphic study of pre-quaternary deposits (especially paleocene and eocene) are summarised here. although the surface morphology of denmark is well mapped and described in detail, there is still room for new observations and reinterpretation of some glacially formed features like rock-cored drumlins from bornholm that are described for the first time. understanding methane distribution in young marine sediments is important for long-term use and protection of areas such as the baltic sea. geus has been partner in the baltic gas project with a key role to map the occurrence of free gas, to quantify the flux of gas and to study the processes and parameters governing methane generation and consumption. natura 2000 is a network of nature protection areas under the habitats directive of the european union. some results from mapping of a natura 2000 area with bubbling reefs in the inner danish waters are presented in this volume. another paper describes how øresund became a strait between 8000 and 9000 years ago due to sea-level rise after an earlier history of being dry land with some bogs and lakes, and later a fjord. the coastal cliff stevns klint is a classical locality for field trips, especially for studying the world’s best exposed cretaceous–tertiary boundary, but also for the beautiful scenery. for this reason an understanding of the risks of cliff collapse is crucial, and an analysis of the different collapse types including volumetric considerations based on photogrammetric studies is given in one of the papers. with a strong danish ambition to reduce the dependence on fossil fuel by using renewable energy sources, shallow geothermal energy can be economically competitive. one paper introduces this concept with comments on exploitation, modelling and groundwater protection. the latter is very important in denmark where almost all drinking water comes from groundwater. the danish policy on this topic and the role of the authorities and water supply companies are discussed. geus is also involved in long-term monitor© 2012 geus. geological survey of denmark and greenland bulletin 26, 7–8. open access: www.geus.dk/publications/bull 88 ing of groundwater above the gas storage facility at stenlille on sjælland. there has been no evidence of leakage through the geological barriers, and only one short incident with traces of gas from an injection well. earthquake activity is low in an intra-plate region like denmark, but prediction of seismicity is important, and the question whether geological and geodetic observations should be included in the assessment is discussed. in 2011 there was a high level of field activities in greenland with a major mapping and geochemical programme in south-east greenland, a large field and shallow-core drilling programme of poorly known cretaceous sediments in north-east greenland and the eager cruise offshore north-east greenland, which was part of the danish continental shelf project. results from these large projects will be presented later; in this issue results from other completed and on-going projects are presented. this volume also contains a discussion of the possibility to use anorthosite from greenland as a future raw material for aluminium instead of bauxite. based on regional knowledge and geochemical data, it is suggested that the fiskenæsset complex is especially interesting. 3d mapping and modelling is high on the agenda in the new geus strategy in both denmark and greenland. different methods are used in the crystalline areas of greenland, for sedimentary basins or for quaternary deposits in denmark. a case study from the skaergaard intrusion in southern east greenland is presented. in greenland, the bureau of minerals and petroleum in nuuk is preparing licensing rounds offshore north-east greenland for 2012 and 2013, and part of the regional preparation for this work is described. there is a high level of exploration activities throughout the west greenland shelf in these years, and one paper brings a case study of bitumen from crystalline rocks in south greenland. climate models predict a marked warming in the arctic that may lead to the release of large amounts of methane bound in permafrost areas or in gas hydrates. the background of and some preliminary results from a scientific cruise to disko bugt in 2011 are related. the important monitoring programme of the greenland ice sheet (promice) continues to develop, and a paper summarises the results of ablation observations from 2008 to 2011. geus takes part in operating many seismic stations in denmark (5) and greenland (18), and it is very time-consuming to analyse all data manually. test of an automatic method is described in another paper, and although the method is efficient, manual processing is still required to detect all earthquakes and for quality control purposes. geus works in many different countries with many different types of projects. the last two papers in this issue are about vietnam and tanzania. geus has been active in vietnam together with our sister organisation vietnam petroleum institute and universities in hanoi for almost twenty years. in tanzania a geocenter denmark project has studied the possibilities of restoring soil fertility by using locally quarried rock phosphate and lime. geological survey of denmark and greenland bulletin 28, 2013, 57-60 57© 2013 geus. geological survey of denmark and greenland bulletin 28, 57–60. open access: www.geus.dk/publications/bull lineament mapping and geological history of the kangerlussuaq region, southern west greenland knud erik s. klint, jon engström, andrew parmenter, timo ruskeeniemi, lillemor claesson liljedahl and anne lehtinen how could future ice ages affect deep nuclear waste repositories in crystalline basement rocks? deep repositories may be affected by a number of glacially induced processes including, but not limited to, (1) fault activation or re-activation and associated seismicity, (2) changing hydraulic and chemical groundwater dynamics and (3) enhanced erosion. such processes are likely to affect not only man-made barriers in spent fuel repositories such as copper canisters and bentonite clay buffers, but also the rock masses that contain and isolate the repositories. in order to increase our understanding of this problem, an international study (the greenland analogue project) was set up in 2008. the aim of the study was to use crystalline bedrock at the margin of the inland ice near kangerlussuaq airport in west greenland as an analogue for future nuclear fuel waste repositories affected by glaciation in fennoscandia and canada. accordingly, a wide range of field surveys were conducted for the analogue project (fig.  1). this paper describes a detailed structural investigation of lineament zones and the establishment of an event succession for fault and fracture zone evolution in central parts of the study area (figs 1b, 2), as well as an interpretation of the distribution of fracture and fault zones with potentially increased permeability. three deep holes were drilled in the study area, and instruments were installed in two of them for subsequent down-hole sampling and monitoring of groundwater to a depth of c. 600 m. the cores were used to compare the subsurface fracture patterns with those established on the basis of surface mapping. fig. 1. a: map of greenland showing the extent of the nagssugtoqidian orogen and the location of fig. 1b (rectangle). b: regional geology of southern west greenland, modified from garde & hollis (2010) and garde & marker (2010). cno, nno and sno are central, northern and southern nagssugtoqidian orogen, respectively. c: lineament map of the area around kangerlussuaq airport. areas a and b were mapped in detail. sisimiut disko bugt lineament systems 1 2 3 5 6 4area a area b (fig. 2)10 km c proterozoic reworked gneiss, mainly granodioritic or quartz dioritic with basic schlieren and lenses of metadolorite dykes orthogneiss late archaean granite early archaean gneiss syntectonic granite suite arfersiorfik igneous suite sisimiut charnockite kangâmiut mafic dykes archaean palaeoproterozoic undifferentiated supracrustal rocks granite amphibolite quaternary pegmatite fault51°w51°w 67°n67°n steep belt shear zo ne ‘flat b elt’northern cno c 52°w a greenland ice sheet 67° 66°n b 50 km southern archaean foreland nagssugtoqidian nno cno sno kangerlussuaqkangerlussuaq nordre isortoq ikertôq thrust zone nordre strømfjord 5858 geological setting the kangerlussuaq area is located close to the southern margin of the c. 1.85 ga old collisional nagssugtoqidian orogen (van gool et al. 2002). the study area covers a 100 × 50 km large area in front of and below the western margin of the inland ice (fig. 1). detailed geological mapping was carried out along a transect from kangerlussuaq airport to the margin of the inland ice, and of a smaller area around the three drilling sites in the valley in front of the inland ice (fig. 2). the bedrock is mainly reworked archaean orthogneiss with minor palaeoproterozoic metavolcanic amphibolite and metasedimentary rocks that were deformed under highgrade metamorphic conditions during the nagssugtoqidian orogeny (van gool et al. 2002). the nagssugtoqidian structures are generally ductile and include a penetrative gneissic fabric, macroscale folds and pronounced shear zones. occasional, deformed mafic dykes also occur, mainly members of the rift-related nagssugtoqidian kangâmiut dyke swarm that preceded the nagssugtoqidian orogeny (mayborn & lesher 2006). brittle structures such as faults and fractures are abundant and were probably formed in a younger, shallower, colder and hence more rigid environment. regional lineament mapping and geology of the study area most lineaments in crystalline rocks represent structural features such as faults and shear zones, rock fabrics and lineaments that were created at discontinuities due to differences in rheology or competence. our lineament mapping comprised four steps. first, lineaments were identified using remotely sensed gis-data compiled from aerial photographs, fig. 2. aerial image of area b (fig. 1c) with mapped and inferred rock fabric (foliation), rock types and structural elements. local event stratigraphic models were developed for locations a–f. the macro-scale structures outline large-scale ductile folds overprinted by various shear and fault zones. three deep holes were drilled (dh-gap01, dh-gap03 and dh-gap04). !! !! !! ! ! (( (( (( (( (( (( (( ( 65° 14 mafic dyke intr f1 f2 kang. dyke intr 74° 60 ° dh-gap01 dh-gap03 dh-gap04 78° 82° décollement f2 folding? dyke intr ( 74 85 8080 79 79 84 84 85 79 86 76 62 70 88 67°08´n kang. dyke intr 1 km fold strike slip fault dip slip fault thrust fault fault foliation fold axis rock fabric banded felsic gneiss mafic gneiss structures syncline hinge line anticline hinge line 1 2 3 4 5 6 kangerlussuaq–russell fau lt lin e lineament system peg intr + 50°09´w lithology: f1 f2 f2 f2 a b c peg intr f2 peg intr peg intr d e peg intr + d f2 f 1 2 3 4 5 6 1 2 3 4 5 6 1 2 3 4 5 61 2 3 4 5 6 1 2 3 4 5 6 1 2 3 4 5 6 1 2 3 4 5 6 russell glacier 59 topographical and geological maps and geophysical data. the second step was a descriptive and kinematic field analysis of the remotely detected lineaments. in the third step, a local event succession model was developed by determination of overprinting relationships. finally, the local geological history of the kangerlussuaq area was interpreted and correlated with the regional geological evolution (van gool et al. 2002; garde & hollis 2010). in order to provide an overview of the geological framework of the study area, the measured and interpreted structures (foliation, fold hingeline, fault and shear zone traces) are shown on a high-resolution aerial image (fig. 2). close to the margin of the inland ice in the north-eastern part of the area, the bedrock is characterised by folded mafic gneiss that defines an open, nnw-trending and shallowly plunging (c. 14°) f1 fold structure. this structure gradually turns into a tight to isoclinal fold that can be traced southward and westward around an open, steeply-plunging, kilometre-scale f2 fold structure with a nwto n-trending axial surface (figs 1c, 2). five general lineament systems were outlined from the gis analysis, and a sixth, subhorizontal system was identified by the field work, during which also the curvilinear nature of the foliation traces became apparent and highlighted the superposition of the f1 and f2 fold phases. on a regional scale, system 1 lineaments are generally distributed between the system 2 steep belts (fig 1). these lineaments dominate in some parts of the southern study area and consist of ductile, e–w-trending, 10–100 m wide, foliation-parallel, steeply inclined and steeply n-dipping dextral shear zones (figs 1c, 2). these shear zones may have acted as décollement planes for the f2 folding (see below). system 2 lineaments constitute the most pronounced linear feature and represent regional, sw-trending shear zones, including the zone outlining søndre strømfjord itself and especially the northern margin of sandflugtdalen that extends from kangerlussuaq to russell glacier, hereafter named the kangerlussuaq–russell fault line. other lineaments farther north with similar orientations (fig. 1b) comprise the ikertôq zone (van der molen 1984), the nordre isortoq steep belt and the nordre strømfjord shear zone (van gool et al. 2002). along the kangerlussuaq–russell fault line this lineament is dominated by sinistral strike-slip movement overprinting elements of thrust faulting dipping towards the north. large boudins of deformed kangâmiut mafic dykes (mayborn & lesher 2006) are included in this zone, which is therefore younger than the intrusion of the kangâmiut dykes. system 3 lineaments are a major nw-trending system along pronounced valleys and extend all the way up to aassiaat and disko bugt (fig. 1a). locally within the study area they form semi-brittle, sinistral strike-slip fault zones that displace the system 1 shear zones. system 4 lineaments form oblique to sub-horizontal fracture systems trending nne and are not represented by any pronounced topographic lineaments. these structures may also be related to a population of normal faults and younger, age general tectonic events in central west greenland correlation to the kangerlussuaq area table 1. event succession of the kangerlussuaq area compared with the general history of central west greenland* * based on van gool et al. (2002) and garde & hollis (2010) † kangerlussuaq–russell fault line >2.5 ga (archaean) formation of ‘banded gneiss’ protoliths f1 folding + multiple healed structures c. 2.04 ga continental rifting coupled with mafic dyke intrusions intrusion of kangâmiut mafic dykes 1.92–1.75 ga nagssugtoqidian orogen. continental collision. reworking, system 1 dextral, strike-slip shear and f2 folding folding and thrusting of gneiss/mafic rocks followed by system 2 wsw–ene-trending faulting (k–r fl†) peak metamorphism and large-scale folding during n–s reactivation of kangerlussuaq–russell fault sinistral contraction. finally formation of steep belts with sinistral strike slip movements system 3 sinistral strike-slip shear zones c. 1.78 to present day various stress conditions. formation and reactivation system 4 normal faults with shallow dip towards of open mode fractures at more shallow depths during se and nw accompanied by pegmatite intrusions 20–25 km uplift 1.2 ga diamond-bearing, ultramafic lamprophyre intrusions s and w of kangerlussuaq 600 ma kimberlite intrusions s and w of kangerlussuaq 100–50 ma faulting related to sea-floor spreading during the system 5 sinistral strike-slip faults trending ne–sw opening of labrador sea and baffin bay. nne–ssw system 6 dextral strike-slip faults trending n–s sinistral strike-slip faulting and conjugate dextral strike slip faulting at least the last 2 ma repeated glaciations resulting in erosion, glacier-induced reactivation of existing fault zones reactivation of fractures during glacial subsidence and rebound of the basement 6060 nne-trending pegmatites orientated parallel with the normal faults. they overprint the system 1 and 3 lineaments, but no cross-cutting relations to system 2 lineaments have been noticed. system 5 lineaments are semi-brittle, sinistral, strike-slip faults trending ne and dipping to the nw. this system outlines the pronounced escarpment in the central to northeastern part of the study area (fig. 2). local zones of wswstriking thrust faults with both northerly and southerly dips are regarded as representing local transpression in a wrench fault system related to the overall system 5 lineaments. system 6 lineaments consist of n–s-trending, brittle, dextral, strike-slip faults. this system may be conjugate with system 5. lineaments with this orientation near the outer coast of central west greenland have been related to the rifting between greenland and canada during the last 100 ma (wilson et al. 2006). preliminary geological history as demonstrated by its structural complexity (fig. 2), the study area has undergone several episodes of deformation. these episodes have been compiled into a local event history that recognises seven types of structural overprint (table 1). two stages of folding (f1 and f2) are identified. f1 may be of archaean age, while f2 seems to be nagssugtoqidian and contemporary with the system 1 shear zones, since the kangâmiut dykes were folded during this event. the kangerlussuaq–russell fault line of system 2 lineaments has been reactivated with sinistral, strike-slip movement, during which the kangâmiut dykes were deformed, but a clear, cross-cutting relation with systems 1 and 3 has not yet been documented. the system 3 lineaments clearly overprint system 1 and the f2 folds. the normal faulting and nne trending pegmatites of system 4 either indicate a general, extensional stress regime or local transtension during strikeslip movements. the timing of this event is highly speculative. the youngest geological events are related to the type 5 and 6 lineaments that form two, generally brittle, strike-slip fault systems. these are also regarded as the primary hydraulic zones in the kangerlussuaq area. the absolute ages of the different types of ductile and brittle deformation are uncertain because of lack of radiometric ages. however, in relative terms, it is suggested that the kangerlussuaq–russell fault line and most other semi-ductile shear zones are ancient features related to the nagssugtoqidian orogeny (van gool et al. 2002). the system 3 and 4 semi-brittle lineaments may be of intermediate ages related to postorogenic tectonic events, whereas the youngest and most brittle deformations may be related to the opening of the north atlantic ocean, the labrador sea and the baffin bay over the last 100 ma (wilson et al. 2006). acknowledgement the study was funded by the swedish, finnish and canadian nuclear waste management organisations. references garde, a.a. & hollis, j.a. 2010: a buried palaeoproterozoic spreading ridge in the northern nagssugtoqidian orogen, west greenland. geological society special publications (london) 338, 213–234. garde, a.a. & marker, m. 2010: geological map of greenland, 1:500 000, søndre strømfjord – nuussuaq. copenhagen: geological survey of denmark and greenland. mayborn, k.r. & lesher, c.e. 2006: origin and evolution of the kangâmiut mafic dyke swarm, west greenland. in: garde, a.a. & kals beek, f. (eds): precambrian crustal evolution and cretaceous–palaeogene faulting in west greenland. geological survey of denmark and greenland bulletin 11, 61–86. van gool, j.a.m., connelly, j.n., marker, m. & mengel, f.c. 2002: the nagssugtoqidian orogen of west greenland: tectonic evolution and regional correlations from a west greenland perspective. canadian journal of earth sciences 39, 665–686. van der molen, i. 1984: dykes and deformation in the ikertôq zone of the nagssugtoqidian at søndre strømfjord airport, west greenland. bulletin of the geological society of denmark 32, 101–106. wilson, r.w., klint, k.e.s., van gool, j.a.m., mccaffrey, k.j.w., holdsworth, r.e. & chalmers, j.a. 2006: faults and fractures in central west greenland: onshore expression of continental break-up and seafloor spreading in the labrador – baffin bay sea. geological survey of denmark and greenland bulletin 11, 185–204. authors’ addresses k.e.s.k., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: kesk@geus.dk j.e. & t.r., geological survey of finland, p.o. box 96, sf-02151 espoo, finland. l.c.l., swedish nuclear fuel and waste management co, box 250, se-101 24, stockholm, sweden. a.p., nuclear waste management organization, 22 st. clair avenue east, sixth floor, toronto, m4t 2s3, canada. a.l., posiva oy olkiluoto, sf-27160 eurajoki, finland. mailto:kesk@geus.dk review article christiansen et al. 2020: geus bulletin 44. 4567. https://doi.org/10.34194/geusb.v44.4567 1 of 21 a review of oil and gas seepage in the nuussuaq basin, west greenland – implications for petroleum exploration flemming g. christiansen*1 , jørgen a. bojesen-koefoed1 , gregers dam1 , troels laier2, sara salehi1 1geological survey of denmark and greenland (geus), copenhagen, denmark, 2geological survey of denmark and greenland (geus), copenhagen, denmark (emeritus) abstract the nuussuaq basin in west greenland has an obvious exploration potential. most of the critical elements are well documented, including structures that could form traps, reservoir rocks, seals and oil and gas seepage that documents petroleum generation. and yet, we still lack a full understanding of the petroleum systems, especially the distribution of mature source rocks in the subsurface and the vertical and lateral migration of petroleum into traps. a recently proposed anticlinal structural model could be very interesting for exploration if evidence of source rocks and migration pathways can be found. in this paper, we review all existing, mostly unpublished, data on gas observations from nuussuaq. furthermore, we present new oil and gas seepage data from the vicinity of the anticline. occurrence of gas within a few kilometres on both sides of the mapped anticline has a strong thermogenic fingerprint, suggesting an origin from oil-prone source rocks with a relatively low thermal maturity. petroleum was extracted from an oil-stained hyaloclastite sample collected in the aaffarsuaq valley in 2019, close to the anticline. biomarker analyses revealed the oil to be a variety of the previously characterised “niaqornaarsuk type,” reported to be formed from campanian-age source rocks. our new analysis places the “niaqornaarsuk type” 10 km from previously documented occurrences and further supports the existence of campanian age deposits developed in source rock facies in the region. 1 introduction the exploration potential for petroleum in the nuussuaq basin in west greenland (figs 1 and 2) was first realised in the early 1990s, based on the observations of oil seepage followed by core drilling and conventional exploration drilling (christiansen 1993, 2011; christiansen et al. 1994a, 1994b, 1995b, 1996a, 1996b, 1997a). however, we currently lack a full understanding of the petroleum systems of the area. although oil seeps have been widely recognised on many coastal localities and classified in detail analytically (bojesen-koefoed et al. 1999, 2007; christiansen et al. 1996c), we still do not know the areal distribution of mature petroleum source rocks in the subsurface and the vertical and lateral migration of oil and gas into possible traps or to the surface. most recently, sørensen et al. (2017) proposed a new play concept based on the photogrammetric mapping of inversion structures. a newly mapped large structural anticline on central nuussuaq (fig. 2) with expected good *correspondence: fgc@geus.dk received: 24 mar 2020 accepted: 08 july 2020 published: 04 dec 2020 keywords: nuussuaq basin, west greenland, gas observations, oil and gas geochemistry, anticline petroleum exploration model abbreviations: dgu: geological survey of denmark ggu: geological survey of greenland geus: geological survey of denmark and greenland gc–ms: gas chromatography–mass spectrometry gcfid: gas chromatography–flame ionization detection mplc: medium-pressure liquid chromatography gc–mssim: selective ion monitoring gc–ms dinsar: differential synthetic aperture radar interferometry ndvi: normalized difference vegetation index geus bulletin is an open access, peerreviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: catherine jex (geus, denmark) reviewed by: chris parry (ger x as, now university of stavanger, norway) and sverre ohm (university of stavanger, norway) funding: see page 19 competing interests: none declared additional files: see page 19 https://doi.org/10.34194/geusb.v44.4567 https://orcid.org/0000-0001-6098-9402 https://orcid.org/0000-0001-5647-2769 https://orcid.org/0000-0002-2905-3136 https://orcid.org/0000-0002-8999-603x mailto:fgc@geus.dk christiansen et al. 2020: geus bulletin 44. 4567. https://doi.org/10.34194/geusb.v44.4567 2 of 21 www.geusbul let in.org reservoirs and seals is an obvious exploration target if mature source rocks and migration pathways can be demonstrated. to understand this target and to provide the necessary input for a risk assessment, we need to further document oil and gas seepage in the inland areas, especially along and east of the kuugannguaq–qunnilik fault zone in central nuussuaq (fig. 2). compared to the numerous oil seeps along the coasts of disko and nuussuaq, only a few records have been obtained inland, where exploration logistics are more complicated and costs are higher. also, freshly eroded rocks along the coast seem to better preserve oil than inland exposures, where volcanic rocks weather differently due to frequent melting and freezing processes and often alter to rocks with a distinct smell of soil. less attention has been paid to document the occurrences of gas in the nuussuaq basin, although some preliminary data were obtained during systematic analyses of boreholes drilled by the geological survey of greenland (ggu), which later merged with the geological survey of denmark to form the geological survey of denmark and greenland (geus) and industry. unfortunately, in a number of these records, gas was not sampled and documented properly or at all. this paper presents a systematic review of all gas observations and data in the nuussuaq basin in order to understand the petroleum systems and aid future exploration. here, we (1) summarise existing data for gas accumulations, much of which were, until now, only available in unpublished ggu and geus reports and (2) present new critical data based on samples collected in 2019. these new data are important to characterise oil and gas seepage near to the kuugannguaq–qunnillik fault zone, where future drilling is being considered. fig. 1 simplified geological map of the nuussuaq basin, west greenland, showing the position of the outcropping sediments on disko, nuussuaq and svartenhuk halvø. ebf: eastern boundary fault. location of umiivik-1 core is indicated. for the purpose of this paper, nuussuaq basin refers to the area shown in this figure. nn nuussuaq umiivik-1 71°n 72°n 72°n 51°w53°w 71°n 70°n 69°n 55°w 51°w svartenhuk halvø ubekendt ejland greenland ice sheet uummannaq innerit hareøen fig. 2 iti lli fa ul t vaigat disko disko bugt ilulissat qeqertarsuaq aasiaat ebf ebf 50 km greenland neogene sediment cover offshore saqqaata qaqqaa central complex paleocene picrites (vaigat formation) undifferentiated basalts offshore naqerloq formation svartenhuk formation maligât formation maastrichtian– paleocene sediments albian–campanian sediments extensional fault precambrian basement fault with lateral or alternating displacements https://doi.org/10.34194/geusb.v44.4567 http://www.geusbulletin.org christiansen et al. 2020: geus bulletin 44. 4567. https://doi.org/10.34194/geusb.v44.4567 3 of 21 www.geusbul let in.org 2 geological setting and exploration models of the nuussuaq basin the nuussuaq basin is a rift basin that developed during cretaceous – paleocene due to extension between canada and greenland. the basin is characterised by outcropping sediments on disko, nuussuaq and svartenhuk halvø. the southern limit is located around qeqertarsuaq, disko (fig. 1), but the northern and western limits are less well defined. for the purposes of this paper, nuussuaq basin corresponds to the area shown in fig. 1. the sediments of the nuussuaq basin and the overlying volcanic rocks are well exposed throughout the disko–nuussuaq–svartenhuk halvø region and are important to understand the sedimentology, stratigraphy, depositional and subsidence history of the sedimentary basins in west greenland. the nuussuaq basin has been intensively studied as an analogue for offshore basins. most of these studies are based on large field campaigns in 1991–1997 and 2004, combined with many shorter field trips (christiansen 1993; christiansen & pulvertaft 1994; christiansen et al. 1992, 1995a, 1996a, 1997a, 1998). for an overview of previous research and exploration history, see dam et al. (2009) and christiansen (2011). the present paper focuses on the parts of the nuussuaq basin on western and central nuussuaq between the itilli and the kuugannguaq–qunnilik fault zones (fig. 2). in this part of the nuussuaq basin, the sedimentary succession is covered by a few kilometres of volcanics of the vaigat and maligât formations (figs 1–3; pedersen et al. 2017, 2018). the underlying sediments were mainly characterised by drilling or from field work on exposures in the itilli valley, along the itilli fault zone (figs 1 and 2). the known sedimentary succession is dominated by marine sediments of mid-cretaceous to palaeogene age. a general facies change occurs from deltaic and slope sediments close to the kuugannguaq–qunnilik fault zone to deeper marine turbidites in and around the itilli fault zone (dam et al. 2009). many mudstone intervals occur in the itilli and kangilia formations (dam et al. 2009); some of these may represent possible source rocks and many are likely to have good sealing capacities (fig. 3). the most likely reservoir intervals are turbidite sandstones in the itilli formation and incised valley sandstones of the quikavsak and agatdal formations and their equivalents (dam & sønderholm 1994, 1998; dam et al. 2009; hjuler et al. 2017; kierkegaard fig. 2 simplified geological map of northern disko and western nuussuaq in the nuussuaq basin. location of wells and cores with oil and gas, major and minor oils seeps, localities with various types of gas observations and the approximate position of the anticline from sørensen et al. (2017; red-dashed line) are marked. 54°30´ gane#1 ganw#1 gank#1 gant#1 gro#3 55° 55°30´ 55°30´ 54°30´ 54°30´55° 70°30´ 70°30´ 70°15´ 70°15´ 70°45´ 54° 54° max o ccurence of a naanaa m b iti lli f au lt zo ne g as sø f au lt zo ne annertuneq core with gas hydrates pingo 132 with gas visman boreholes with gas aaffarsuaq valley serfat cores with gas k uu ga nn gu aq –q un ni lik f au lt zo ne v a i g a t m a l i g â t fault zone lake, pingo with gas major oil seep minor oil seep outcrop of basement outcrop of volcanic rocks outcrop of cretaceous-paleocene sediments covered core hole with gas core hole with oil core hole with oil and gas deep well with oil and gas sikillinge asuk 10 km marraat-1 new pingo gassø lake gassø lake tr ac e of a nt ic lin e 574305 a gatdalen iti lli va lle y https://doi.org/10.34194/geusb.v44.4567 http://www.geusbulletin.org christiansen et al. 2020: geus bulletin 44. 4567. https://doi.org/10.34194/geusb.v44.4567 4 of 21 www.geusbul let in.org fig. 3 simplified sedimentary and volcanic stratigraphy of the nuussuaq basin on disko, nuussuaq and svartenhuk halvø. the relationship between petroleum seeps and stains and main petroleum systems elements is shown. possible ages for source rocks of the following oil types are shown: m: marraat type; n: niaqornaarsuk type; i: itilli type; k: kuugannguaq type. the eqalulik type source rock is not known, but it is assumed to be early cretaceous. many of the petroleum stains in the vaigat formation are associated with later eocene dykes. vaigat formation includes the following: an: anaanaa member; na: naujánguit member; or: ordlingassoq member. note that the eocene volcanic units are rather thin and not regionally distributed. hareøen formation is only found on hareøen, erqua formation only on ubekendt ejland and naqerloq formation only on hareøen and westernmost parts of nuussuaq, ubekendt ejland and svartenhuk halvø. based on sørensen et al. (2017), dam et al. (2009) and pedersen et al. (2017, 2018). upper cretaceous paleocene eocene oligocene miocene pliocene quaternary lower cretaceous 0 ma 5 10 15 20 25 30 35 40 45 50 55 60 65 70 75 80 85 90 95 100 105 110 115 120 aptian albian cenomanian turonian coniacian santonian campanian maastrichtian danian selandian thanetian ypresian lutetian bartonian priabonian rupelian chattian aquitanian burdigalian langhian serravallian tortonian mainly continental deposits mainly marine or deltaic sandy/silty deposits, locally mudstones mainly marine mudstones, locally sandy/silty messinian zanclean piacenzian p al eo ce ne e oc en e 61 –6 0 60 –5 8 56 –5 4 54 –5 3 39 –3 8 62 –6 1 m al ig aâ t fm s va rt en hu k fm n aq er lo q f m e rq ua fm h ar eø en fm va ig at f m ma volcanic stratigraphy chronostratigraphy sedimentary stratigraphy nw se nw se source rock unconformity (hareøen) (hareøen) kangilia fm itilli fm itilli fm atane fm agatdal fm eqalulik fm m n k i quikavsak fm lake, pingo with gas core hole with gas core hole with oil core hole with oil and gas major oil seep minor oil stain an or na https://doi.org/10.34194/geusb.v44.4567 http://www.geusbulletin.org christiansen et al. 2020: geus bulletin 44. 4567. https://doi.org/10.34194/geusb.v44.4567 5 of 21 www.geusbul let in.org 1998). the most likely traps are extensional-rotated fault blocks formed in the cretaceous or palaeogene, or late paleocene inversion structures such as the anticline mapped by sørensen et al. (2017). nuussuaq experienced a short, but very intensive, exploration phase in 1994–1998, driven largely by new, and at that time unpublished, data on oil seepage. the small canadian company grønarctic energy inc. managed to drill four fully cored stratigraphic boreholes (ganw#1, gane#1, gank#1, and gant#1) and one deep wildcat well (gro#3) in 1994–1996. but, despite many encouraging oil shows and documentation of good reservoirs and seals in the upper cretaceous and palaeogene succession, the company was unable to raise funding for further drilling and they eventually relinquished their licences in may 1998. 3 existing data from oil seeps and bore holes in the nuussuaq basin the first oil-seeps on nuussuaq were discovered in the marraat area in 1992 (christiansen 1993). since then, significant time and resources have been invested to find additional localities, especially along the coast, with evidence of either visible oil seepage or micro-seepage in mineralised veins. oil seepage is very common in an area on western nuussuaq (christiansen et al. 1996c). following the years of systematic field work, evidence of oil has also been found in many other localities, including disko, hareøen, ubekendt ejland and svartenhuk halvø (figs 1 and 2). several hundred samples of oil seeps or oil-impregnated cores on nuussuaq were analysed in detail by bojesen-koefoed et al. (1997a, 1999), and supplemented more recently by bojesen-koefoed et al. (2007). they described the characteristics of the oil samples from the region and grouped them into five distinct oil types. these types are adopted here, but for a complete overview of the oil-type classification, we refer the reader to bojesen-koefoed et al. (1999, 2007). the main models of source rock distribution, depositional environment and generative history of the source rocks are based on state-of-the-art analyses and are documented in detail (bojesen-koefoed et al. 1999, 2004, 2007; christiansen et al. 1996c). some of the oils have a unique composition of biological markers, containing organic compounds such as lupanes and a series of norhopanes that were only rarely documented at the time of analyses (nytoft et al. 2000, 2002). understanding the distribution and concentration of different oil types is important for petroleum exploration on nuussuaq. previous studies demonstrated working petroleum systems and first indications of where the source rocks could be expected in the subsurface, where and when they have generated oil and how migration, and in some cases degradation, took place. the oils so far recorded in the nuussuaq basin occur in two main settings: 1. in oil-impregnated porous lavas and hyaloclastites that may have formed exhumed continuous reservoirs in the deeper part of the volcanic succession, especially within the anaanaa member of the vaigat formation (pedersen et al. 1998) or just below, in the uppermost part of the sedimentary successions (fig. 3). oil-impregnated rocks hold large volumes of hydrocarbons, which were generated, and probably migrated vertically, from an underlying source rock such as oil of the “marraat deltaic type” or the lesser known “eqalulik type” or “niaqornaarsuk type” (as defined by bojesen-koefoed et al. 1999). this migration likely occurred during and shortly after the main phase of volcanism in the region (62–60 ma) with rapid subsidence and possibly increased heat flow. such oils occur over large parts of western nuussuaq, especially in the area from marraat-1–gane#1– sikillinge (fig. 2), where several billion barrels of more or less degraded oil may fill out most available porosity in the volcanics (christiansen et al. 2006; see supplementary file s6). 2. in migration conduits, especially along faults, dykes, fractures or as fluid inclusions in thin mineralised veins in many different volcanic units (fig. 3). oils also occur in some sands in the asuk area on disko (fig. 2). such oils are generally low in volume and concentration, but are known from large areas on disko, nuussuaq, ubekendt ejland and svartenhuk halvø, where they often belong to the “itilli type,” presumed to be generated from a marine mid-cretaceous source rock (fig. 3; bojesen-koefoed et al. 1999, 2007). the first setting offers some possibilities for local exploration, especially in incised valley deposits of the lower paleocene quikavsak and agatdal formations on western nuussuaq. the area of exploration interest is, however, rather small with complex structural features that are not likely to define large targets. the second setting suggests good exploration opportunities over much larger areas in the deeper part of the cretaceous succession and supports the anticlinal model suggested by sørensen et al. (2017). the main area of known oil seeps and stains was significantly enlarged after numerous field seasons. it is likely that the area containing oil of the “marraat type” can be further extended inland towards the north and northeast. the easternmost record of the “marraat type” is the gank#1 borehole. it is unclear whether this oil https://doi.org/10.34194/geusb.v44.4567 http://www.geusbulletin.org christiansen et al. 2020: geus bulletin 44. 4567. https://doi.org/10.34194/geusb.v44.4567 6 of 21 www.geusbul let in.org type can be traced further into the kuugannguaq–qunnillik fault zone (fig. 2). the main challenge for future oil seep studies is to find more examples from the second setting, especially along a possible fairway from northern nuussuaq along the anticline to central and southern nuussuaq. the major fault zones in the region are also interesting targets for future studies, as they may have been important migration pathways. 4 existing records of gas in the nuussuaq basin well-documented analytical records of gas observations in the nuussuaq basin are relatively few. there are some historical observations of gas leakages from lakes or pingos, but they have not been analysed with modern analytical methods. more recent records were obtained during drilling campaigns, which revealed a high likelihood of widely distributed gas under pressure in the subsurface of the nuussuaq basin, either below a permafrost seal or deeper in the sedimentary succession. since 1992, improved sampling techniques were developed for drilling and field work by the ggu and later by geus, which ensured a better understanding of gas distribution in the subsurface. here, we review all known records of gas observations in chronological order and describe the various sampling techniques used, their limitations, to provide key results and make preliminary interpretations. 4.1 numerous ggu boreholes, drilled in 1992 numerous cores were drilled at agatdalen (fig. 2), annertuneq (fig. 2) and svartenhuk halvø (same location as umiivik-1; fig. 1) in 1992. these shallow fully cored boreholes were drilled to depths between 45 and 95 m by ggu using a custom-made light-weight rig that could be mobilised using small helicopters. the main goal was to document the presence of oil-prone source rocks within the marine cretaceous succession (fig. 3; christiansen 1993; christiansen et al. 1994c). core samples for gas analysis were taken in metal tubes. results were reported by laier (1994) in an unpublished institutional report, which is provided here as supplementary file s1. gas amounts were relatively low, which precluded stable isotopic analysis. methane concentrations were low relative to ethane, propane and butane, which suggest preferential leakage of lighter molecules. for this reason, cans were used in subsequent studies, replacing the tubes. during the drilling at annertuneq (core number 400407), white to bluish gas hydrates were observed at a depth of c. 7 m (fig. 4). at the time of drilling, it was not realised that the material was gas hydrates and samples evaporated before they could be properly described and secured for analysis. 4.2 marraat-1 core, drilled in 1993 marraat-1 (408001) core (location fig. 2; sample numbers 408011 and 408020 in table 1) was drilled by the canadian company falconbridge ltd. for ggu in august 1993. it terminated at a depth of 448 m. the well was subsequently logged in october–november 1993 and some additional fluid samples of formation water were taken. the main goal was to clarify if the solid bitumen found at the surface was an indication of undegraded oil in the subsurface (christiansen et al. 1994a; dam & christiansen 1994). the subsequent analytical programme focused on the oil composition that suggested new models for age and depositional environment of the source rock (bojesen-koefoed et al. 1999; christiansen et al. 1994b, 1996c). some core pieces were sealed in cans for subsequent gas and formation fluid analyses. these data were reported by laier (1994; see supplementary file s1) and christiansen et al. (1995b) and are summarised in table 1. two samples of a rather dry gas had a sufficient concentration of methane to allow analysis of stable carbon isotopes. values of δ 13c (13c/12c) and wetness indicate a mixed thermogenic biogenic origin (table 1; fig. 5). 4.3 falconbridge mineral exploration cores, drilled in 1994 during their mineral exploration programme for nickel sulphides in 1994, falconbridge ltd. observed gas bubbles and froth on core surfaces. they penetrated a zone of pressured gas at c. 290 m depth in one of the boreholes in the serfat area (core number fp94-11-04; fig. 2). the gas was found in cretaceous sediments below thick sills on the north coast of nuussuaq (dam & nøhr-hansen 1995). five samples of gas were obtained from the cores and stored in plastic containers. data were provided fig. 4 bluish gas hydrates at c. 7 m depth in core 400407 at annertuneq on the north coast of nuussuaq (location in fig. 2). core diameter is 3.0 cm. photo taken on 1 august 1992. https://doi.org/10.34194/geusb.v44.4567 http://www.geusbulletin.org christiansen et al. 2020: geus bulletin 44. 4567. https://doi.org/10.34194/geusb.v44.4567 7 of 21 www.geusbul let in.org ta bl e 1  g as c om po si tio ns fo r se ve n co re s in th e n uu ss ua q ba si n, w es t g re en la nd co re (c or e nu m be r) sa m pl e nu m be r or c or e pi ec e d ep th (m ) ch 4 (p pm ) c 2h 6 (p pm ) c 3h 8 (p pm ) ic 4h 10 (p pm ) nc 4h 10 (p pm ) ic 5h 12 (p pm ) nc 5h 12 (p pm ) δ13 c 1 (‰ ) δ13 c 2 (‰ ) δ13 c 3 (‰ ) δd c 1 (‰ ) w et ne ss g as o ri gi n d at a so ur ce s m ar ra at -1 (4 08 00 1) 40 80 11 1 41 .0 28 .9 0. 08 −5 3. 4 36 1 th la ie r 19 94 ; s up pl em en ta ry fi le s 1 (a ls o in cl ud e ad di tio na l a na ly se s) 40 80 20 1 82 .0 37 .1 0. 10 −5 3. 4 37 1 th se rf at (f p9 411 -0 4) 38 00 02 2 29 0. 0 10 42 36 1 13 6 3. 1 8. 8 0. 5 0. 5 −1 1. 2 −2 0. 3 −2 0. 9 2. 1 th h m la ie r un pu bl is he d da ta o ct ob er an d d ec em be r 19 94 38 00 03 29 0. 2 22 52 51 6 13 1 3. 1 10 .2 1. 0 1. 8 3. 5 su pp le m en ta ry fi le s s2 a nd s 3 38 00 04 29 0. 4 27 65 49 3 13 5 3. 4 10 .4 0. 8 1. 4 4. 4 38 00 05 29 0. 6 23 06 54 1 16 8 4. 7 14 .0 1. 0 1. 5 3. 6 38 00 06 2 29 0. 8 20 06 48 1 15 5 4. 8 13 .0 0. 9 1. 5 −2 2. 2 −2 3. 7 −2 1. 9 3. 1 th h m g an w #1 (3 80 10 1) 38 01 05 3 72 1. 0 83 2 00 0 61 0 00 13 0 00 15 00 10 00 −4 3. 3 −2 8. 1 −2 7. 1 −1 99 .0 11 2 th l m ch ri st ia ns en e t a l. 19 95 b g an e# 1 (4 39 00 1) 43 90 07 3 63 1. 0 82 3 00 0 28 8 −4 5. 4 22 0 00 th l m ch ri st ia ns en e t a l. 19 96 b (a ls o in cl ud e ad di tio na l a na ly se s) 43 90 01 -3 68 63 3. 7 44 4 00 33 3 33 0 67 .3 14 4 28 18 −4 5. 1 67 th l m 43 90 01 -5 30 68 9. 8 91 00 4 −4 0. 9 22 60 th m m g an k# 1 (4 39 20 1) 43 92 01 -0 54 19 4. 5 46 7 00 21 10 38 4 54 30 .8 10 .4 3. 5 −4 9. 5 18 .7 th l m ch ri st ia ns en e t a l. 19 96 b (a ls o in cl ud e ad di tio na l a na ly se s) 43 92 01 -1 13 36 5. 9 23 0 00 0 53 00 19 20 12 2 15 2 11 7 −4 6. 9 −3 2. 3 31 .9 th l m 43 92 01 -1 18 37 9. 7 13 3 00 0 46 30 15 40 14 0 12 2 24 8 −4 1. 9 21 .6 th l m g an t# 1 (4 39 10 1) 43 91 07 3 24 7. 2 76 00 −6 5. 4 0 bi og en ic ch ri st ia ns en e t a l. 19 96 b 43 91 11 3 60 8. 8 26 9 00 24 0 30 −3 4. 8 10 0 th h m 43 91 12 3 60 8. 8 64 2 00 0 41 0 −3 5. 0 15 66 th h m 43 91 16 60 8. 8 37 2 00 36 16 .3 1. 2 3. 9 0. 5 1 −3 7. 0 71 1 th h m 43 91 01 -4 20 64 9. 4 18 8 90 0 33 50 56 2 68 19 2 30 18 −4 6. 0 −3 0. 4 48 .2 th l m 43 91 01 -4 47 73 7. 0 82 0 00 25 6 40 .2   −4 0. 1 27 7 th m m 43 91 01 -4 49 74 3. 7 26 4 00 0 39 0 16 00 86 24 2 28 39 −4 0. 4 13 3 th m m 43 91 01 -4 57 2 76 9. 3 23 0 00 15 60 27 8 8. 8 21 .7 1. 9 2. 6 −3 6. 8 12 .5 th h m 43 91 01 -4 58 77 4. 0 10 7 00 0 19 10 34 9 3 7 n. a. n. a. −3 9. 7 47 .1 th m m 43 91 01 -4 65 79 4. 0 43 2 00 63 2 10 1 4. 1 10 .9 0. 9 1. 3 −3 9. 1 58 .9 th m m 43 91 01 -4 72 2 81 6. 2 38 8 00 19 3 82 5. 7 10 .3 0. 9 1. 3 −2 0. 8 14 1 th h m 43 91 28 3 90 1. 3 33 9 00 −6 8. 0 0 bi og en ic co nt in ue d https://doi.org/10.34194/geusb.v44.4567 http://www.geusbulletin.org christiansen et al. 2020: geus bulletin 44. 4567. https://doi.org/10.34194/geusb.v44.4567 8 of 21 www.geusbul let in.org ta bl e 1  g as c om po si tio ns fo r se ve n co re s in th e n uu ss ua q ba si n, w es t g re en la nd (c on tin ue d) co re (c or e nu m be r) sa m pl e nu m be r or c or e pi ec e d ep th (m ) ch 4 c 2h 6 c 3h 8 ic 4h 10 nc 4h 10 ic 5h 12 nc 5h 12 δ13 c 1 (‰ ) δ13 c 2 (‰ ) δ13 c 3 (‰ ) δd c 1 (‰ ) w et ne ss g as o ri gi n d at a so ur ce s u m iiv ik -1 (4 39 30 1) 10 4. 7 33 1 00 0 72 8 79 7 22 3 15 1 −5 1. 2 21 7 th l m ch ri st ia ns en e t a l. 19 97 b (a ll an al ys es ) 15 1. 3 44 9 00 26 50 78 2 21 5 84 34 .7 −2 1. 1 13 .1 th l m 19 9. 7 20 1 40 0 20 20 46 4 tr tr −4 3. 7 81 .1 th l m 25 0. 7 39 4 80 17 50 59 9 39 .5 56 .7 −2 0. 6 16 .8 th l m d am e t a l. 19 98 (a dd iti on al in te rp re ta tio n) 30 7. 7 14 8 00 0 20 17 66 0 tr tr −3 8. 9 55 .3 th l m 40 3. 6 77 5 00 25 70 58 0 40 92 −2 4. 6 24 .6 th h m 45 0. 5 68 5 00 37 00 50 0 78 82 −2 1. 9 16 .3 th p m 75 2. 5 40 4 20 0 17 6 00 19 40 15 5 17 5 −3 1. 8 20 .7 th p m 79 4. 0 43 9 20 64 00 12 80 16 7 12 4 −2 7. 6 5. 7 th p m 91 0. 4 41 0 96 0 51 0 80 65 40 59 0 29 3 29 .6 −3 2. 6 −2 7. 9 −2 4. 4 7. 1 th p m 10 65 .6 49 9 40 55 40 92 5 10 0 10 0 −2 0. 1 7. 7 th p m 11 51 .4 42 5 00 0 77 1 20 12 9 30 24 70 14 90 65 8 18 4 −3 5. 1 −2 7. 6 −2 3. 6 4. 7 th p m 11 63 .6 34 8 00 0 79 5 00 17 4 60 40 00 21 60 87 2 21 9 −3 4. 6 −2 7. 8 −2 4. 6 3. 6 th p m 11 71 .9 41 0 00 0 46 8 00 93 40 20 00 14 00 62 2 20 0 −3 2. 9 −2 7. 6 −2 3. 5 7. 3 th p m 11 82 .6 38 9 60 92 10 29 10 69 0 57 0 25 0 92 −1 1. 4 −2 0. 7 −2 1. 5 3. 2 th p m 11 97 .9 31 3 00 0 21 00 12 70 27 0 25 8 95 41 .5 −3 7. 8 −2 2. 2 −2 4. 2 92 .8 th p m 1 co nt ai ne d hi gh n itr og en . 2 d iff er en tia l l ea ka ge o f c 1 a nd li gh t i so to pe s. 3 s te el c yl in de r. tr : t ra ce a m ou nt s. b la nk c el ls in di ca te th at th e co m po un d is no t p re se nt (b el ow d et ec tio n lim it) o r t ha t i so to pe s w er e no t a na ly se d. c h 4 (c 1): m et ha ne ; c 2h 6 (c 2): et ha ne ; c 3h 8 (c 3): pr op an e; ic 4h 10 : i so bu ta ne ; n c 4h 10 : n -b ut an e; ic 5h 12 : i so pe nt an e; n c 5h 12 : n -p en ta ne ; w et ne ss : c 1/( c 2 + c 3); th : t he rm og en ic ; l m : l ow -t he rm al m at ur ity w ith re sp ec t t o oi l g en er ati on ; m m : m ed iu m th er m al m at ur ity w ith re sp ec t t o oi l g en er ati on ; p m : p os tm at ur e w ith re sp ec t t o oi l g en er ati on ; n .a .: no t a na ly se d. https://doi.org/10.34194/geusb.v44.4567 http://www.geusbulletin.org christiansen et al. 2020: geus bulletin 44. 4567. https://doi.org/10.34194/geusb.v44.4567 9 of 21 www.geusbul let in.org by laier in october and december 1994 as unpublished data (see supplementary files s2 and s3, respectively) and are summarised in table 1. the samples contained significant volumes of gas with relatively high concentrations (up to 3000 ppm) of wet gases. c-isotope composition of methane (δ13c1), ethane (δ13c2) and propane (δ13c3) suggested a thermogenic origin from a source rock with a relatively high thermal maturity. the data suggested a loss of lighter isotopes by diffusion from the plastic containers (table 1). this was tested by experiments of the containers (laier, unpublished data, december 1994; supplementary file s3). 4.4 ganw#1, drilled in 1994 ganw#1 (core 380101, fig. 2) was drilled by grønarctic energy, calgary, canada in september–october 1994 as a follow-up to the marraat-1 borehole (christiansen et al. 1995a). the main goal was to penetrate the base of the volcanic succession and to document further oil impregnation at depth. one gas sample (380105) from a depth of 721 m was sampled from the wellhead in a steel cylinder, and a full suite of gas analyses was carried out. the data were reported by christiansen et al. (1995b) and are presented in table 1. the gas had a moderate wetness (table 1; fig. 5). δ13c1 versus δd of methane (δdc1) suggests a thermogenic origin from a source rock and an association with oil (fig. 6), while δ13c1 versus δ13c2 suggests a low thermal maturity dominated by type iii kerogen (fig. 7; christiansen et al. 1995b). 4.5 gane#1, gank#1 and gant#1, drilled in 1995 three fully cored boreholes gane#1 (core 439001), gank#1 (core 439201) and gant#1 (core 439101) and one sidetrack (gane#1a; data not shown here) were drilled by grønarctic energy in the summer of 1995 to depths of between 398 and 901 m as part of their exploration and production license on western nuussuaq (locations in fig. 2). the main goal was to characterise the sedimentary succession below the volcanic succession and to demonstrate an active petroleum system. all boreholes revealed oil and gas within volcanic or sedimentary rocks. detailed sedimentological and stratigraphical studies and comprehensive geochemical analyses of organic compounds were carried out by ggu for grønarctic energy (christiansen et al. 1996b). some of these data are presented in table 1. some gases sampled in steel cylinders together with gas from core-pieces sealed in cans were analysed. gas was commonly observed in many intervals in gant#1 fig. 5 wetness (c1/c2+c3) versus δ13c of methane (δ13c1) for marraat-1, ganw#1, umiivik-1 and pingo 132. c1: methane; c2: ethane; c3: propane. compositional fields indicate biogenic or thermogenic origin. modified from schoell (1984). 10000 1000 100 10 1 c 1/ (c 2 + c 3) –40 –60 –80 bacterial thermogenic ganw#1 umiivik-1 pingo 132marraat-1 δ13c1 fig. 6 δ13c of methane (δ13c1) versus δd of methane (δdc1) for ganw#1 and pingo 132. plotted compositional fields (blue lines) are from jenden and kaplan (1989). –300 –200 –100 δdc1 –20 –30 –40 –50 –60 –70 microbial gas mixed gas oil-associated gas condensateassociated gas δ1 3 c 1 ganw#1 pingo 132 https://doi.org/10.34194/geusb.v44.4567 http://www.geusbulletin.org christiansen et al. 2020: geus bulletin 44. 4567. https://doi.org/10.34194/geusb.v44.4567 10 of 21 www.geusbul let in.org and gane#1, indicated by bubbles in the drilling fluids. most notable was the gas flaring of gane#1 at a depth of c. 660 m (see fig. 4 in christiansen et al. 1996a). corresponding cores from this drilling depth show oil impregnation with the relatively rare “eqalulik type” that cannot be correlated to any known source rock (bojesen-koefoed et al. 1997a, 1999). in gane#1 (and gane#1a), gas was commonly observed in several sandstone intervals (631–641, 684– 689 and 696–702 m). in gant#1, gas was commonly observed in many sandstone intervals between 575 and 775 m. most of these gases are thermogenic in origin. but their variable composition suggests the presence of both low maturity gases from the penetrated succession and high maturity gases that may have migrated from deeper in the subsurface (table 1). 4.6 umiivik-1, drilled in 1995 umiivik-1 (core 439301) was drilled as a 1200 m deep stratigraphic well by grønarctic energy for ggu in august to september 1995. the main goal was to test and document a cenomanian–turonian source rock (bate & christiansen 1996; dam et al. 1998). in the deeper part of the well, gas was heard to be escaping the core. some intervals revealed a white froth on the core surface when it was removed from the core barrel (bate & christiansen 1996). twenty-seven core pieces were sealed in cans and analysed for their gas composition (christiansen et al. 1997b; dam et al. 1998). sixteen of these are presented in table 1 – the remaining nine samples had no detectable amounts of gas. gas concentrations of the 16 samples were high, with significant amounts of wet gases such as propane, butane and pentane (table 1). in some deeper parts of the well, concentrations were so high that the sampling cans deformed. compositions in the deeper part are typical of thermogenic gas associated with oil (table 1; figs 5 and 7). unfortunately, the isotopic composition trend suggests some diffusion after sampling as suggested by christiansen et al. (1997b). the presence of a postmature oil-prone source rock in the deeper part (below 1100 m) of umiivik-1 was documented in more detail by drits et al. (2007). 4.7 gro#3, drilled in 1996 gro#3 was drilled by grønarctic energy in the summer of 1996, following promising results from previous drilling and seep studies (christiansen et al. 1997a, 1998). cores or sidewall cores were not included in the drilling programme, and the organic geochemical results are based on analysis of cuttings only (bojesen-koefoed et al. 1997b; christiansen et al. 1998). eight sandstone intervals were drillstem tested to obtain fluid samples, but results were inconclusive. later, log interpretation fig. 7 δ13c of methane (δ13c1) versus δ13c of ethane (δ13c2) for ganw#1, umiivik-1 and pingo 132. maturity lines are calculated from faber (1987) for type ii kerogen. ro: vitrinite reflectance. –40 –35 –30 –25 –20 –50 –45 –40 –35 –30 δ1 3 c m et ha ne 0.5 %ro 0.7 %ro 0.9 %ro 1.1 %ro 1.3 %ro 1.5 %ro 1.8 %ro 2.0 %ro 2.5 %ro 3.0 %romix, different thermogenic gases or microbial methane oxydation mix, microbial methane ganw-1 umiivik-1 pingo 132 δ13c ethane + + + + + + + + + + https://doi.org/10.34194/geusb.v44.4567 http://www.geusbulletin.org christiansen et al. 2020: geus bulletin 44. 4567. https://doi.org/10.34194/geusb.v44.4567 11 of 21 www.geusbul let in.org indicated many intervals with high gas concentrations (kristensen & dam 1997). these petrophysical data are not presented in this review. 4.8 vismann mineral exploration, drilled in 2007 the company vismann exploration inc. drilled two mineral exploration boreholes in the aaffarsuaq valley, in 2007, based on previously observed geophysical anomalies in the area (fig. 2). the logistical operation was complex and required construction of a new road into the aaffarsuaq valley. both of the wells were suspended due to gas under pressure at depths of 154 and 133 m. neither of the holes reached bedrock and they only penetrated the quaternary overburden (glacial tills). unfortunately, no gas was sampled. 4.9 lakes and pingos the best opportunity to observe gas seepage in land terrains, like the nuussuaq basin, is from lakes, pingos, below newly formed ice or on partly wet mud flats. pingos are conspicuous mound-like landforms that are common in regions with continuous permafrost. they may have craters resembling those of mud volcanoes (pissart 1988). active pingos are formed by periglacial processes, have an ice core of frozen water and often grow over time. some eventually collapse. pingos are common in many valleys on disko, nuussuaq and svartenhuk halvø. expedition anecdotes of bubbling lakes suggested the presence of gas seepage as early as the 1930s. early analyses of both gas and water collected in the 1930s and 1940s were first presented in the context of petroleum exploration by henderson (1969). these early analyses document a significant content of methane and alkaline water associated with gas seepage. many of the pingos in the region were more systematically studied in the 1990s – most of them were dry. some occasionally show crater lakes or outlets of spring water under artesian pressure. sampling of these waters may give information on the composition of water and gases below the permafrost. pingo 132 (fig. 2) north of the aaffarsuaq river seems to have been rather active and wet over many decades. note that henderson (1969) uses the term qapiortoq kitdleq for the same pingo. pingo 132 was visited and sampled on several occasions in 1991 and 1992. snow fans were observed to disappear from pingo 132 later than on any other southward facing slopes in this part of the aaffarsuaq valley. on one occasion, a fountain of water under pressure was observed (fig. 8a). a similar feature was documented in a photo taken on 25 august 1939 by b. thomsen (see fig. 6 in henderson 1969). looking downstream from the snow fan, which covers part of the pingo, the valley floor is described as being overgrown with algae and other vegetation, suggesting that the outlet has been active over long periods of the year and that the water is rich in nutrients. this remarkable colouration may be a good proxy for remote-sensing studies of other similar outlets in the region. elsewhere, gas seepage from the hinterland of marraat is indicated by the so-called gassø lake (fig. 2), depicted in the 1:100 000 geological map of the area (rosenkrantz et al. 1974) and the official geodætisk institut 1:250 000 topographical map from 1980. the lake was visited by flemming g. christiansen and inger salomonsen on 28 july 1994. the lake surface showed clear evidence of seepage as intense bubbling (fig. 8b). analyses of lake water (sampled in cans) indicated a complete dominance of nitrogen (laier, unpublished data, october 1994; supplementary file s2), suggesting either long-distance migration of gases that are thermally very mature, or more likely that oxygen had been lost by bacteria in an anoxic environment. a third and more recent example is from the marraat area, where a new pingo seems to be actively forming (figs 2 and 8c). the normally flat riverbed surface is doming with new fractures in the peaty soil. the underlying ice-core is beginning to be exposed and large bubbles of gas are visible in small ponds nearby, beneath recently formed ice (fig. 8d). such features with fractures opening to permafrost below – and possibly with degrading permafrost – may become more common on nuussuaq in the years to come, especially in areas with active movements. 4.10 marine records geophysical data, including conventional seismic data and high-resolution shallow seismic data, indicate that gas could be very common in the sedimentary succession below the seabed of vaigat (location in fig. 2). examples of so-called direct hydrocarbon indicators have been observed offshore in vaigat, particularly as flat spots but also as gas cloud features (bojesen-koefoed et al. 2007). geochemistry data of pore waters from gravity cores indicate that gas hydrates may also be present at several places offshore in disko bugt and vaigat (mikkelsen et al. 2012; nielsen et al. 2014). this is supported by numerous observations of pockmarks, seabed mud diapirs and change in reflection patterns on geophysical data in the area. kuijpers et al. (2001) also observed intense degassing from two cores south of disko. to the best of our knowledge, no gas samples have been analysed. 4.11 summary: existing gas observations historical observations of gas seepage in the nuussuaq basin are to some degree supported by modern analytical data. large parts of the nuussuaq basin are clearly https://doi.org/10.34194/geusb.v44.4567 http://www.geusbulletin.org christiansen et al. 2020: geus bulletin 44. 4567. https://doi.org/10.34194/geusb.v44.4567 12 of 21 www.geusbul let in.org b c d a fig. 8 features of pingo 132 in aaffarsuaq valley, gassø lake, and a new pingo near marraat-1 core. a: water fountain indicating water with a high gas content under artesian pressure below the permafrost, 17 august 1991. b: surface of gassø lake with clear indications of gas seepage, 28 july 1994. c: doming, soil fracturing and possible formation of a new pingo near marraat-1, 27 july 2006 (photo: roy fitzsimmons). d: gas trapped under ice in a small pool near marraat-1, 27 july 2006. https://doi.org/10.34194/geusb.v44.4567 http://www.geusbulletin.org christiansen et al. 2020: geus bulletin 44. 4567. https://doi.org/10.34194/geusb.v44.4567 13 of 21 www.geusbul let in.org underlain by a sedimentary succession that contains high concentrations of gas. although dominated by biogenic gas, gases from surface lakes and pingos show a distinct thermogenic component. the gases from boreholes are mainly thermogenic in origin, and in some cases, their composition suggests an association with oil. these oil-associated gases have a rather low thermal maturity corresponding to the thermal maturity of the sediments penetrated by drilling. in several cases, examples of thermally highmaturity gases have been recorded. these gases may have migrated from the deeper part of the sedimentary succession or were generated in the vicinity of dykes and sills. clearly, better systematic sampling techniques, proper handling and storage, and most importantly, rapid analyses using modern instrumentation could provide much more valuable information. gant#1, pingo 132 and the vismann mineral exploration holes are all in close proximity to the anticline, suggested by sørensen et al. (2017) to be a large potential target for future exploration. it is therefore particularly important to get more data on oil and gas seeps from this area. 5 2019 reconnaissance to sample oil and gas seepage 5.1 biomarkers in oil seepage near the kuugannguaq–qunnilik fault zone on 26 july 2019, we visited a number of planned drill sites next to the kuugannguaq–qunnilik fault zone in the aaffarsuaq valley to check for oil seepage (fig. 2). hyaloclastites from the deep part of the vaigat formation were examined for signs of petroleum staining. the hyaloclastites are from unit 409 of the nuusap qaqqarsua member within the naujánguit member of the vaigat formation (fig. 3; see details in pedersen et al. 2002, 2017). we sampled the hyaloclastites, located c. 340 m.a.s.l., close to the outlet of the qunnillik canyon (sample site 574305 in fig. 2), a few hundred metres west of the expected trace of the kuugannguaq– qunnillik fault zone. we picked out small pieces of a hard, fresh rock with thin carbonate veins and a distinct petroliferous odour for organic geochemical analyses using standard methods (bojesen-koefoed et al. 2018). a sample of rock pieces was lightly crushed and extracted for 4 h (1 h immersed in boiling solvent followed by 3 h of rinsing) using a soxhtec™ instrument and a 93+7 vol./vol. dichloromethane + methanol mixture as solvent. the extract was recovered by evaporation over n2 and weighed. a 238 g sample was extracted to obtain a total yield of 6.2 mg extract, corresponding to c. 26 ppm. asphaltenes were precipitated by the addition of 40-fold excess n-pentane. asphaltenes were recovered by centrifugation and rinsed through several stages with n-pentane. asphaltenes account for 27.4% by weight of the total. maltene (i.e. asphaltene-free) fractions were separated in saturated hydrocarbons, aromatic hydrocarbons and polar fractions by medium-pressure liquid chromatography using a procedure modified from radke et al. (1980). the maltene fraction is dominated by polar nso compounds (table 2). the saturated hydrocarbon fraction was analysed by gas chromatography–flame ionization detection (gcfid) using a shimadzu gas chromatograph, furnished with a 30-m wcot zb-1 capillary column. biomarker analysis was carried out by gas chromatography–mass spectrometry (gc–ms) using an agilent 6890n gas chromatograph, fitted with a 30 m wcot zb-5 capillary column, coupled to a waters (micromass) quattro micro gc tandem quadrupole–hexapole–quadrupole ms. the instrument was run in both selective ion monitoring mode (gc–mssim) and gc–ms–ms parent–daughter mode. the  sample was run several times using methods designed to optimise the representation of different compounds. gas chromatographic data on the saturated hydrocarbon fraction show a strongly front-end evaporated distribution of n-alkanes, a high proportion of longchain components (fig. 9a) and no unresolved complex mixture, suggesting limited biodegradation (fig. 9a). front-end losses of short-chain components make calculation of standard ratios futile, including the pristane/ phytane ratio. the concentrations of tricylic terpanes (fig. 9b) are relatively low, and their distribution partially obscured by the presence of abundant unknown components in the same range, probably various other triand tetracyclic components. pentacyclic triterpanes show a series of hopanes ranging from c27 to c35, including notable proportions of 28.30-bisnorhopane (h28, fig. 9c) and oleanane (o, fig. 9c), plus trace amounts of bicadinanes and taraxastane (not shown). extended 28-bisnorhopanes are absent, as are nor/bisnorlupanes. the bishomohopane isomerisation ratio has reached equilibrium table 2 maltene fraction extracted from oil-stained hyaloclastics (sample site 574305, fig. 2) in the affarsuaq valley, nuussuaq basin sample saturated hydrocarbons (wt.%) aromatic hydrocarbons (wt.%) polar compounds (wt.%) 574305 10.0 2.5 87.5 nso: nitrogen, sulphur, oxygen. https://doi.org/10.34194/geusb.v44.4567 http://www.geusbulletin.org christiansen et al. 2020: geus bulletin 44. 4567. https://doi.org/10.34194/geusb.v44.4567 14 of 21 www.geusbul let in.org at 0.61 (h32 s/(s+r), table 3, fig. 9c). the sample shows high concentrations of aromatic diterpanes which are a common feature of terrestrial oils found on nuussuaq, moderate concentrations of phenanthrene and methylphenanthrene (not shown), and the presence of diand triaromatic oleanane. the sterane distribution is strongly dominated by c29 moieties and shows very high proportions of diasteranes (fig.  9d). c30 desmethyl steranes are absent, but c26-steranes are relatively prominent and allow calculation of a nordiacholestane ratio of 0.37 (holba et al. 1998). c29-sterane 20s/(20s+20r) and αββ/ (ααα+αββ) isomerisation ratios are both below equilibrium at 0.43 and 0.42, respectively (table 3, fig. 9d). sample 574305 can be classified as a “niaqornarsuk type” oil with some notable deviations, according to its biological-marker characteristics as first defined by bojesen-koefoed et al. (1999; table 4). a few parameters fall outside the established range for this oil type, notably the relative abundance of diasteranes and 28,30-bisnorhopane. however, the following diagnostic criteria are fulfilled: 1. appreciable concentrations of 28,30-bisnorhopane and absence of extended 28-norhopanes (table 4). 2. presence of oleanane (perhaps coeluting with small amounts of lupane, see nytoft et al. [2020]), with no or negligible concentrations of nor/bisnorlupanes (table 4, fig. 9c). 3. strong predominance of c29-steranes and absence of c30 desmethyl steranes (table 4, fig. 9d). the aromatic fingerprint, in particular the moderate concentrations of phenanthrene and methylphenanthrene, further supports the identification of sample 574305 as “niaqornarsuk type” oil. the characteristic features of the “niaqornarsuk type” oil were originally defined using only gc–ms(sim) data, which are inferior to modern gc–ms–ms. sterane data based on gc–ms(sim) suffer from coelution problems, which often cause misleadingly low ratios of diasteranes to regular steranes. the niaqornaarsuk oil type has been linked to campanian-age source rocks, based on the geochemical correlation to campanian-age shales of the gant#1 borehole (bojesen-koefoed et al. 1999), which are perfectly conformable with a nordiacholestane ratio of 0.37 (holba et al. 1998). the sample was collected at least 10 km from any other known occurrence of surface seepage near the kuugannguaq–qunnillik fault zone, which is encouraging for future exploration. the presence of a niaqornaarsuk oil type at a considerable distance from the only hitherto known occurrences of this oil type further supports the presumed existence of campanian age deposits developed in source-rock facies in the region. gcfid gc-mssim m/z 191 gc-ms-ms pentacyclics sum of 9 transitions gc-ms-ms steranes sum of 5 transitions p ris ta ne p hy ta ne nc 20 nc 25 nc 30 nc 35 h 30 h 29 h 28 o + l tm ts 29 ts m 29 m 30 h 31 (s + r ) h 32 (s + r ) h 33 (s + r ) s29 αα αs αα αr αβ βs αβ βrd 27 d 27 d 27d 27 a b c d fig. 9 gcfid and gc–ms–ms data (time vs. signal) for a sample of oil-stained hyaloclastics (sample site 574305; fig. 2), aaffarsuaq valley, nuussuaq basin. a: gcfid data. b: gc–mssim for m/z 191. c: gc– ms–ms of pentacyclics. sum of nine transitions. d: gc–ms–ms of steranes. sum of five transitions. symbols are as follows: ncx: normal alkanes (x = carbon number); pristane: c19 acyclic isoprenoid; phytane: c20 acyclic isoprenoid; ts: 18α-trisnorneohopane; tm: 17α-trisnorhopane; h28: 28,30-bisnorhopane; h29: norhopane; 29ts: c29 neohopane; m29: normoretane; o+l: coelution of oleanane + lupane; h30: hopane; m30: moretane; hxx (s+r): homohopanes, doublets 22s and 22r isomers; xx = carbon number; d27: c27 diasteranes; s29: c29 regular steranes; αααs: regular sterane ααα 20s isomer; αββs: regular sterane αββ 20s isomer; αββr: regular sterane αββ 20r isomer; αααr: regular sterane ααα 20r isomer. https://doi.org/10.34194/geusb.v44.4567 http://www.geusbulletin.org christiansen et al. 2020: geus bulletin 44. 4567. https://doi.org/10.34194/geusb.v44.4567 15 of 21 www.geusbul let in.org 5.2 gas seepage in aaffarsuaq valley 5.2.1 remote-sensing analysis of pingo 132 on 26 july 2019, pingo 132 in the aaffarsuaq valley, central nuussuaq, was visited to check for mud extrusion and gas seepage (fig. 2). although pingo 132 is periglacial in origin, it resembles typical mud volcanos from classical petroliferous basins or geothermal fields with mud overflow (see etiope 2015; mazzini & etiope 2017; mazzini et al. 2011). today, there is evidence of recent mud extrusion in the area. the mud cropping out is pale in colour, has rather steep sides and many irregular fractures and erosional features that are not likely to survive more than a few winters. there is little to no vegetation on the extruded mudstone, which is in contrast to the older parts of the pingo and surrounding valley floor. the actual mudstone contains numerous clasts of rounded basements boulders, some cretaceous atane formation sandstone lithologies and a few volcanic rock types – all typical of the quaternary tills in the aaffarsuaq valley (fig. 10). high-resolution satellite images provide good possibilities for detecting surface movements of the pingo and measuring slow displacement rates of centimetres to metres per year. using differential synthetic aperture radar (sar) interferometry (dinsar; rosen et al. 2000) to plot the phase differences between two or more satellite sar images allows us to detect movement in the direction of line-of-sight of a few millimetres and helps characterise the dynamics of terrain uplift. combining the two complementary techniques overcomes the limitations of using just one of these methods. to visually identify the changes in the shape of pingo 132 and outcropping mud over time, we used optical four-band planetscope time-series with 5 m spatial resolution between 2017 and 2019 (fig. 11a). the normalized difference vegetation index (ndvi) is calculated from the same dataset (fig. 11b) to enhance the table 3 key geochemistry parameters for sample 574305 sample c19–26 tricyclics/ hopane c23 tricyclic/ hopane c25/c26 tricyclics c24 tetracylic/ hopane h32 s/(s+r) ts/(ts+tm) s29 s/(s+r) s29 αββ/ (ααα+αββ) 574305 0.09 0.01 0.83 0.01 0.61 0.41 0.43 0.42 h32 s/(s+r): bishomohopane isomerisation ratio; ts/(ts + tm): 18α-trisnorneohopane/(18α-trisnorheohopane + 17α-trisnorhopane); s29 s/(s+r): c29 sterane 20s/(20s+20r) isomerisation ratio; s29 αββ/(ααα+αββ): c29 sterane αββ/(ααα+αββ) isomerisation ratio. table 4 comparison of sample 574305 with the niaqornaarsuk oil type (bojesen-koefoed et al. 1999) sample h28/h29 h29/h30 o/(o + h30) d27/rs27 rs27/rs29 s27% s28% s29% 574305 0.31 1.14 0.18 4.0 0.32 19 21 60 niaqornaarsuk oil type mean 0.17 0.92 0.06 0.87 0.32 20 14 65 minimum 0.13 0.81 0.04 0.73 0.20 15 12 57 maximum 0.23 1.02 0.10 1.19 0.47 27 17 73 h28/h29: 28,30-bisnorhopane to norhopane ratio; h29/h30: norhopane to hopane ratio; o/(o+h30): oleanane to oleanane + hopane ratio d27/ rs27: c27 diasterane to c27 regular sterane ratio; rs27/rs29: c27 to c29 ratio of regular steranes; s27%, s28%, s29%: relative distribution of regular steranes. fig. 10 extruding muds from the side of pingo 132. note the many rounded basements boulders in the mud. height of section c. 3 m. photo taken on 26 july 2019. https://doi.org/10.34194/geusb.v44.4567 http://www.geusbulletin.org christiansen et al. 2020: geus bulletin 44. 4567. https://doi.org/10.34194/geusb.v44.4567 16 of 21 www.geusbul let in.org 200 m a b c july 2017 july 2018 july 2019 august 2019august 2018august 2017 18 august 2017 14 august 2019 10–22 august 2019 5 august 201813 july 2017 12 july 2019 5–17 july 2019 11–23 june 2019 200 m 200 m -3 -2 -1 -1 1 0 1 2 3 fig. 11 remote-sensing images and analysis of pingo 132, affarsuaqq valley. red-dashed lines indicate the location of pingo 132. a: optical four-band planetscope data with 5 m spatial resolution for july and august 2017–2019. b: normalized vegetation index (ndvi) for july and august 2017 and 2019. negative values of ndvi correspond to water. values close to zero (–0.1 to 0.1) generally correspond to barren areas of rock, sand or snow. low positive values (0.2–0.4) represent shrubs, while high values (approaching 1) indicate green vegetation. c:close-up of pingo 132. selected differential interferograms (wrapped interferometric phase) from track 175. upper row: 10 august 2019–22 august 2019 (temporal baseline: 12 days; normal baseline: –0.06 m). lower row: 5 july 2019–17 july 2019 (temporal baseline: 6 days; normal baseline: 15.68 m) and 11 june 2019–23 june 2019 (temporal baseline: 6 days; normal baseline: 5.46 m). https://doi.org/10.34194/geusb.v44.4567 http://www.geusbulletin.org christiansen et al. 2020: geus bulletin 44. 4567. https://doi.org/10.34194/geusb.v44.4567 17 of 21 www.geusbul let in.org interpretation of the identified patterns and to measure the concentration of green vegetation. two sentinel-1 tracks (i.e. 171 descending and track 90 ascending) cover the same area. differential sar interferometry was carried out for the descending track between 11 june and 3 september 2019 using the arctic digital elevation model (porter et al. 2018). the interferograms were unwrapped, and the deformational rates are reported as the satellite line-of-sight rates projected onto the steepest slope. the results indicate significant vertical movements in both summer and winter, suggesting “uplift” rates of c. 1 cm every 12 days, probably related to mud accumulation (fig. 11c). furthermore, seasonal variation in the uplift rate seen in the dinsar data matches the seasonal pattern observed in the optical data. it seems that most of the observed mud extrusion took place in the summer of 2017. 5.2.2 geochemistry of gas seepage the present water outlet is on the lower, south side of the pingo. it is associated with fractures in the soil and peat and small ponds with bubbling gas (see videos in supplementary files s4 and s5). some of the partly dry mudflats show gas-escape vents (fig. 12a). the sandier material displays a crater-like feature, a few centimetres in size (fig. 12b). a gas sample was taken in a plastic bottle where gas displaced the outlet water, kept cool and analysed within a week for ch4 and c2h6 by standard gas chromatography (christiansen et al. 1997b). the sample was stored and later analysed for stable carbon and hydrogen isotopes by martin krüger at bundesanstalt für geowissenschaften und rohstoffe in hannover (for methods, see blumenberg et al. 2016). the seeping gas is mainly composed of methane with a small amount of ethane (table 5). the carbon isotope composition of methane and ethane using standard classification plots suggests a thermogenic origin with a relatively low thermal maturity (figs 5–7). 5.2.3 geochemistry of water associated with gas accumulation or seepage geochemistry of formation water associated with oil and gas accumulations or related to oil and gas seepage may provide important additional information on the migration and degradation history. water under pressure has been recorded in a few places on nuussuaq, both in the marraat-1 and gank#1 wells and in some pingos. some historical data were published by henderson (1969), and additional data from the early nineties were compiled and reported by christiansen et al. (1995b). these are presented in table 5 along with new data for pingo 132. the formation fluids from marraat-1 have a higher salinity than seawater and a very high ca/mg ratio suggesting a deep brine origin (table 6). there is some variation between different levels, suggesting that the fig. 12 gas seepage from pingo 132. a: gas escape structures on recently dried-out mud flat on the southern side of pingo 132, 26 july 2019. size is c. 20 × 30 cm. b: centimetre-scale mounds formed by gas seepage at pingo 132, 26 july 2019. hammerhead for scale. table 5 geochemistry of gas escaping from pingo 132 sampled in 1991, 1992 and 2019 sample number date ch4 (c1) (ppm) c2h6 (c2) (ppm) δ13c1 (‰) δ13c2 (‰) δdc1 (‰) wetness data sources 358472 17 august 1991 723 000 530 −45.8 n.a. n.a. 1364 laier 1994 400843 14 july 1992 468 000 n.d. −40.4 n.a. n.a. n.d. laier 1994 400844 14 july 1992 81 800 790 −38.1 n.a. n.a. 1035 laier 1994 400894 14 august 1992 355 000 175 −43.4 n.a. n.a. 2028 laier 1994 547303 26 july 2019 239 000 164 −43.2 −34.5 −233 1460 this study n.a.: not analysed; n.d.: not determined. wetness: c1/(c2 + c3). c3 not present in any samples. https://doi.org/10.34194/geusb.v44.4567 http://www.geusbulletin.org christiansen et al. 2020: geus bulletin 44. 4567. https://doi.org/10.34194/geusb.v44.4567 18 of 21 www.geusbul let in.org different volcanic lithologies and their content of zeolites could affect composition. pingo 132 is less saline than marraat-1 (table 6). it should be noted that in pingo 132, there is an increased salinity compared to river water, and with high na/k ratios, low ca/mg ratios and low cl/so4 ratios (table 6). the 1991 and 1992 samples (fountain water only) have a rather consistent composition through time and a slightly enriched ph between 8.78 and 8.97 (average: 8.89) compared to the river, sea and formation waters and a high alkalinity (table 6). the 2019 sample was collected in a small pool and seems to be dominated by surface water from melting snow. 6 implications for exploration and recommendation for future studies the 2019 and previously documented gas data and 2019 oil seep data from nuussuaq support an exploration model for the anticlinal structures mapped by sørensen et al. (2017). petroleum extracted from an oil-stained hyaloclastite sample, collected in the aaffarsuaq valley, in 2019 represents a facies variety of the “niaqornaarsuk type” sensu bojesen-koefoed et al. (1999). the presence of a “niaqornaarsuk type” oil 10 km from other known occurrences of this oil type further supports the presumed existence of campanian age deposits developed in source-rock facies in the region. importantly, we observed no sign of mixing with the “marraat type,” suggesting that the marraat source rock disappears somewhere between gank#1 and the kuugannguaq– qunnilik fault zone, or that the source rock, if present, is thermally immature. furthermore, numerous examples of gas occur within a few kilometres on either side of the mapped anticline and along possible migration pathways. these gases have a thermogenic fingerprint and suggest a possible origin from oil-prone source rocks with a relatively low thermal maturity. further geological and structural mapping using 3d photogrammetry combined with geophysical data would be an ideal approach to develop the exploration model in the region. moreover, future studies should table 6 geochemistry of water samples from pingo 132 and nearby formation water, seawater and river water sample location type ph alk cl (mg/l) so4 (mg/l) na (mg/l) k (mg/l) ca (mg/l) mg cl/so4 na/k ca/mg data sources 358472 pingo 132 fountain water 8.78 46.6 256 200 1135 9.48 3.68 86 1.28 119..7 0.04 laier 1994; christiansen et al. 1995b 400843 pingo 132 fountain water 8.97 42.6 248 178 1110 9.74 1.68 68.5 1.39 114.0 0.02 laier 1994; christiansen et al. 1995b 400844 pingo 132 fountain water 8.97 29.7 53.1 177 750 6.72 2.46 54.5 0.30 111.6 0.05 laier 1994; christiansen et al. 1995b 400894 pingo 132 fountain water 8.78 44.8 258 180 1140 9.28 2.63 76.5 1.43 122.8 0.03 laier 1994; christiansen et al. 1995b 547303 pingo 132 fountain water 7.87 3.17 19.9 2.16 69.6 1.52 4.30 5.71 9.2 45.7 0.75 this study 408011 marraat-1 formation water (41 m depth) 7.42 0.92 29 200 1484 9650 188 6740 790 19.68 51.3 8.53 laier 1994; christiansen et al. 1995b 408021 marraat-1 formation water (82 m depth) 7.41 0.58 29 600 1252 8180 126 8560 630 23.64 64.9 13.59 laier 1994; christiansen et al. 1995b 408035 marraat-1 water under pressure (346 m depth) 7.15 0.43 26 500 1276 6780 167 8160 630 20.77 40.6 12.95 laier 1994; christiansen et al. 1995b 408036 vaigat, maraat-1 sea water 7.94 2.10 19 500 2340 10 500 402 430 1230 8.33 25.2 0.36 laier 1994; christiansen et al. 1995b 380132 vaigat, ganw#1 sea water 7.94 2.20 18 930 2033 8797 403 412 1217 9.31 21.8 0.34 laier 1994; christiansen et al. 1995b 380133 ganw#1 river water 8.20 3.20 13 11 20.7 0.29 32.4 17.5 1.18 71.4 1.85 laier 1994; christiansen et al. 1995b alk: alkalinity. https://doi.org/10.34194/geusb.v44.4567 http://www.geusbulletin.org christiansen et al. 2020: geus bulletin 44. 4567. https://doi.org/10.34194/geusb.v44.4567 19 of 21 www.geusbul let in.org systematically sample oil traces along faults and fractures and focus on rock types with carbonate-filled veins that often host fluid inclusions, to elucidate the vertical and lateral distribution of the active petroleum systems in the nuussuaq basin. in some ways, this compares to the practice of traditional onshore exploration in areas like california and texas more than a hundred years ago, where targets were often defined by a combination of surface structures and seeps. this rule of thumb is still valid in many onshore areas around the world, but knowledge of the distinct oil types can guide exploration even more efficiently. systematic mapping, sampling and characterisation of gas seepage from pingos, lakes and thawing permafrost could be similarly important in the future. professional sampling tools for both onshore and offshore activities, including transport and storage of samples under cool conditions, are important, and samples should be analysed as soon as possible to reduce contamination and diffusion. modern isotope techniques with better resolution and low detection limits are likely to provide more details compared to the preliminary work of the 1990s. with degrading permafrost and some specific pingos experiencing rapid change, many more sampling sites are likely to be identified. based on pingo 132, it is obvious that satellite data providing both optimal images and interferograms can systematically identify areas of degrading permafrost. this would allow us to identify suitable sites to collect samples of gas and water that originate from below the permafrost seal. furthermore, satellite data can be used for preliminary dating of mud extrusions. finally, it must be emphasised that the changes observed over the last decades and years point towards a dynamic situation caused by climate change. this may potentially lead to much more frequent mud diapirism and emissions of gas in large parts of the nuussuaq basin when more permafrost degrades in the future. there is a strong need for many of the localities to be documented and monitored in detail. this has implications not only for petroleum exploration but also from a viewpoint of nature preservation as many new local ecosystems are likely to develop and change over time in the coming decades. acknowledgements analytical data are from many different field projects by ggu/geus with additional funding from danish and greenlandic authorities. contributions came especially from the former danish ministry of energy research program (efp), the previous mineral resources administration for greenland in copenhagen, the previous bureau of minerals and petroleum in nuuk and recently from the present ministry of industry, energy, research and labour, nuuk, greenland. close collaboration with industry over the years, especially grønarctic energy inc. and its enthusiastic manager cam hanna has been very useful. practical help over many years and analytical support by john boserup, ditte kiel-duhring, h. p. nytoft and christina rosenberg lynge are highly appreciated. martin krüger at bundesanstalt für geowissenschaften und rohstoffe in hannover is gratefully acknowledged for performing isotopic analyses of the pingo gas samples collected in 2019. we would like to thank dietmar backes from the university of luxembourg for giving us access to planetscope imagery provided via the planet’s education and research program that has been used for generation of ndvi maps. comments on the paper by the reviewers chris parry and sverre e. ohm were welcome. additional information funding statement samples from 2019 were collected on a field trip financed by ministry of industry, energy, research and labour, nuuk, greenland. other costs were paid by the geological survey of denmark and greenland. author contribution fgc: writing the original draft (lead). jabk: analyses, presentation and discussion of oil geochemistry data. gd: contribution to historical data and petroleum exploration model. tl: analyses, presentation and discussion of water and gas geochemistry data. ss: analyses, presentation and discussion of satellite data. additional files six additional files are available online: 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(eds): arctic petroleum geology. geological society (london) memoir 35, 647–661. https://doi.org/10.1144/m35.42 table 6 geochemistry of water samples from pingo 132 and nearby formation water, seawater and river water sample location type ph alk cl (mg/l) so4 (mg/l) na (mg/l) k (mg/l) ca (mg/l) mg cl/so4 na/k ca/mg data sources 358472 pingo 132 fountain water 8.78 46.6 256 200 1135 9.48 3.68 86 1.28 119..7 0.04 laier 1994; christiansen et al. 1995b 400843 pingo 132 fountain water 8.97 42.6 248 178 1110 9.74 1.68 68.5 1.39 114.0 0.02 laier 1994; christiansen et al. 1995b 400844 pingo 132 fountain water 8.97 29.7 53.1 177 750 6.72 2.46 54.5 0.30 111.6 0.05 laier 1994; christiansen et al. 1995b 400894 pingo 132 fountain water 8.78 44.8 258 180 1140 9.28 2.63 76.5 1.43 122.8 0.03 laier 1994; christiansen et al. 1995b 547303 pingo 132 fountain water 7.87 3.17 19.9 2.16 69.6 1.52 4.30 5.71 9.2 45.7 0.75 this study 408011 marraat-1 formation water (41 m depth) 7.42 0.92 29 200 1484 9650 188 6740 790 19.68 51.3 8.53 laier 1994; christiansen et al. 1995b 408021 marraat-1 formation water (82 m depth) 7.41 0.58 29 600 1252 8180 126 8560 630 23.64 64.9 13.59 laier 1994; christiansen et al. 1995b 408035 marraat-1 water under pressure (346 m depth) 7.15 0.43 26 500 1276 6780 167 8160 630 20.77 40.6 12.95 laier 1994; christiansen et al. 1995b 408036 vaigat, maraat-1 sea water 7.94 2.10 19 500 2340 10 500 402 430 1230 8.33 25.2 0.36 laier 1994; christiansen et al. 1995b 380132 vaigat, ganw#1 sea water 7.94 2.20 18 930 2033 8797 403 412 1217 9.31 21.8 0.34 laier 1994; christiansen et al. 1995b 380133 ganw#1 river water 8.20 3.20 13 11 20.7 0.29 32.4 17.5 1.18 71.4 1.85 laier 1994; christiansen et al. 1995b alk: alkalinity. https://doi.org/10.34194/geusb.v44.4567 http://www.geusbulletin.org https://doi.org/10.22008/fk2/so5vld� https://doi.org/10.22008/fk2/so5vld� https://doi.org/10.1016/j.marpetgeo.2016.05.031 https://doi.org/10.1144/0050305 https://doi.org/10.1144/0050305 https://doi.org/10.34194/geusb.v4.4783 https://doi.org/10.1111/j.1747-5457.2007.00219.x https://doi.org/10.1111/j.1747-5457.2007.00219.x https://doi.org/10.34194/geusb.v42.4314 https://doi.org/10.34194/geusb.v42.4314 https://doi.org/10.1144/m35.42 christiansen et al. 2020: geus bulletin 44. 4567. https://doi.org/10.34194/geusb.v44.4567 20 of 21 www.geusbul let in.org christiansen, f.g. & pulvertaft, t.c.r. 1994: petroleum-geological activities in 1993: oil source rocks the dominant theme of the season’s field programme. grønlands geologiske undersøgelse rapport 160, 52–56. christiansen, f.g. et al.1992: renewed petroleum geological studies onshore west greenland. grønlands geologiske undersøgelse rapport 155, 31–35. christiansen, f.g., dam, g. & pedersen, a.k. 1994a: discovery of live oil at marraat, nuussuaq, west greenland – field work, drilling and logging. grønlands geologiske undersøgelse rapport 160, 57–63. christiansen, f.g., bojesen-koefoed, j.a. & nytoft, h.p. 1994b: organic 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thrane 2004, this volume). the thrust sheets incorporate archaean and palaeoproterozoic orthogneiss complexes overlain by a thick late mesoproterozoic – early neoproterozoic metasedimentary succession known as the krummedal supracrustal sequence; the latter is structurally overlain by the neoproterozoic eleonore bay supergroup and tillite geus bulletin 6.pmd 10-02-2005, 09:5467 68 group, and lower palaeozoic rocks. these rock units have been variably reworked during the caledonian orogeny. samples in this study, pbsl analyses were carried out on garnet and kyanite from samples 426050 and 422766 of the late mesoproterozoic – early neoproterozoic krummedal supracrustal sequence (kalsbeek et al. 2000), with the objective of constraining the age of metamorphism which produced these minerals. in addition, pbsl analyses were undertaken on titanite from a garnet amphibolite (426037) and a gabbroic gneiss (426018), and on an apatite fraction from a tonalitic basement gneiss (426008). the latter three samples all derive from the crystalline basement complex underlying the krummedal supracrustal sequence (fig. 1), and the aim was to determine metamorphic ages for the mineral phases in the samples. whole-rock pb analyses were also carried out on samples 426008, 426018 and 426050. zircon grains from the crystalline basement complex in the study area analysed by ion microprobe have hitherto only yielded archaean and palaeoproterozoic magmatic ages. although the rocks form parts of caledonian thrust sheets and have undergone extensive caledonian deformation and metamorphism, so far not a single caledonian age has been obtained from zircon (thrane 2002). while caledonian k-ar mineral ages have previously been recorded from all rock types in the study area (e.g. rex & higgins 1985), the spread in ages and uncertainties inherent in the method is indicative only of a metamorphic overprint of approximately caledonian age. 27 27º nathorst landcharcot land gletscherland suess land andrée land c a le d o n ia n f o ld b el t archaean mainly gneiss complexes palaeoproterozoic mainly gneiss complexes neoproterozoic (eleonore bay supergroup) to ordovician caledonian granites post-caledonian mesoproterozoic krummedal supracrustal sequence 72º 426008 426018 426050 422766 50 km extensional fault archaean–palaeoproterozoic boundary thrust detachment 456037 ff fig. 1. simplified geological map of the study area in the east greenland caledonides, with sample localities discussed in the text. supracrustal rocks of palaeoproterozoic age in charcot land are included with the palaeoproterozoic gneiss complexes. ff, forsblad fjord. geus bulletin 6.pmd 10-02-2005, 09:5468 69 methods for the samples analysed in the present study, a 200 mm sieve fraction of each mineral separate was purified by hand-picking. the samples were digested in a series of steps using procedures documented in table 1; the method used was modified after that of berger & braun (1997) and frei et al. (1997). the purified pb was loaded on re filaments with silica gel and h 3 po 4, and the isotopic ratios analysed on the vg sector 54it instrument at the university of copenhagen. most analyses were performed using the faraday multi-collector; a few steps that contained very little pb were analysed with the single collector (ion counting daly detector). fractionation of pb was monitored by repeated analyses of the nbs 981 standard (values of todt et al. 1993) and amounted to 0.103 ± 0.016 %/ amu. the calculations of regression lines follow the method of ludwig (1999). errors quoted are 2 f. five to seven acid-leach steps were undertaken on each mineral separate. whole-rock pb and pbsl isotope data are listed in table 2 and plotted in figs 2–5. pbsl results krummedal supracrustal sequence the late mesoproterozoic – early neoproterozoic krummedal supracrustal sequence is widely distributed in the southern part of the east greenland caledonides between 70° and 74°n (fig. 1; higgins 1988). sample 426050 was collected from the krummedal supracrustal sequence south of innermost forsblad fjord, close to the faulted contact with the crystalline basement complexes (fig. 1). the general metamorphic grade of the krummedal supracrustal sequence is amphibolite facies, and the sample consists of quartz + plagioclase + k-feldspar + garnet + biotite + kyanite + sillimanite + amphibole + muscovite + titanite. the garnet and biotite represent early phases, while kyanite is a later phase that overgrows the deformation fabric of the biotite. kyanite and k-feldspar crystallised at the same time, demonstrating that the rock has been exposed to high p–t conditions; during cooling, sillimanite, titanite and secondary biotite crystallised, and part of the kyanite was consumed during formation of muscovite. garnet and kyanite were analysed by pbsl (fig. 2). seven steps were performed on the garnet; all the steps, together with the whole-rock analysis, fall on a linear array in the 207pb/204pb vs. 206pb/204pb diagram table 1. sample data stepwise dissolution procedures mix = 1.5n hbr – 2n hcl 12:1 mixture. all steps except number 1 were left on the hotplate during the dissolution time. 236 mix 10 4.4n hbr 45 8.8n hbr 3 8.8n hbr 24 conc. hf 48 conc. hf 260 8.8n hbr 290 97 mix 10 4.4n hbr 45 8.8n hbr 3 8.8n hbr 24 conc. hf 48 conc. hf 260 719 mix 30 1.0n hbr 60 4.0n hbr 3 8.8n hbr 6 8.8n hbr 12 conc. hf 24 conc. hf 340 195 mix 10 4.4n hbr 45 8.8n hbr 3 8.8n hbr 24 conc. hf 48 conc. hf 260 71 mix 10 4.4n hbr 90 8.8n hbr 3 8.8n hbr 24 conc. hf 50 conc. hf 340 139 10 60 1.0n hbr 3 1.5n hbr 3 8.8n hbr 3 7n hno3 9 sample weight step 1 time step 2 time step 3 time step 4 time step 5 time step 6 time step 7 time mg min. min. hrs hrs hrs hrs hrs 426050 garnet 426050 kyanite 422766 kyanite 426018 titanite 426037 titanite 426008 50% mix + 50% mix + apatite 50% h2o 50% h2o geus bulletin 6.pmd 10-02-2005, 09:5469 70 table 2. pb-pb step leaching (pbsl) data sample phase step 206pb/204pb ± 2 σ* 207pb/204pb ± 2 σ* 208pb/204pb ± 2 σ* r1 r2 426050 wr 19.39 0.01 15.65 0.01 38.91 0.03 0.96 0.93 426008 wr 14.32 0.01 14.61 0.01 38.51 0.03 0.93 0.92 426018 wr 17.10 0.04 15.13 0.04 37.76 0.09 0.99 0.98 426050 grt 1 23.34 0.64 16.01 0.44 39.09 1.08 1.00 1.00 426050 grt 2 29.02 0.75 16.34 0.43 43.66 1.14 0.97 0.98 426050 grt 3 69.99 1.99 18.74 0.53 116.25 3.30 1.00 1.00 426050 grt 4 272.79 6.53 32.93 0.79 436.00 10.43 1.00 1.00 426050 grt 5 99.56 1.31 20.49 0.27 44.90 0.59 1.00 1.00 426050 grt 6 139.41 0.74 23.99 0.13 57.17 0.31 1.00 1.00 426050 grt 7 152.05 4.39 24.48 0.71 87.67 2.53 1.00 1.00 426050 ky 1 19.53 0.06 15.66 0.05 38.17 0.11 0.99 0.98 426050 ky 2 20.45 0.14 15.70 0.11 38.52 0.27 0.99 1.00 426050 ky 3 24.26 0.23 15.82 0.15 47.05 0.44 0.99 0.99 426050 ky 4 34.83 0.16 16.42 0.08 65.68 0.31 0.99 0.99 426050 ky 5 20.00 0.02 15.70 0.02 38.09 0.05 0.95 0.94 426050 ky 6 33.13 0.19 16.49 0.10 38.81 0.22 0.99 0.99 422766 ky 1 18.23 0.05 15.49 0.04 37.96 0.10 0.99 0.98 422766 ky 2 19.92 0.16 15.64 0.12 40.81 0.32 0.99 1.00 422766 ky 3 42.22 0.52 16.75 0.21 82.30 1.03 0.97 0.98 422766 ky 4 171.43 19.09 24.12 2.69 326.96 36.41 1.00 1.00 422766 ky 5 137.46 5.87 22.08 0.95 258.40 11.04 1.00 1.00 422766 ky 6 20.59 0.04 15.70 0.03 38.60 0.07 0.98 0.97 422766 ky 7 27.19 1.02 16.16 0.61 39.21 1.48 1.00 1.00 426018 tit 1 16.50 0.02 15.09 0.02 37.40 0.04 0.95 0.95 426018 tit 2 19.47 0.05 15.27 0.04 38.22 0.11 0.99 0.99 426018 tit 3 104.74 1.13 20.29 0.22 56.20 0.61 1.00 1.00 426018 tit 4 152.44 0.53 22.89 0.08 65.29 0.23 0.99 0.99 426018 tit 5 120.05 0.35 20.97 0.06 56.99 0.17 0.99 0.99 426018 tit 6 122.84 0.92 21.14 0.16 57.72 0.43 1.00 1.00 426037 tit 1 16.53 0.22 15.11 0.20 37.11 0.50 1.00 1.00 426037 tit 2 20.87 0.09 15.30 0.07 40.03 0.18 0.99 0.99 426037 tit 3 76.66 1.92 18.21 0.46 46.56 1.17 1.00 1.00 426037 tit 4 69.70 1.09 17.90 0.28 48.30 0.76 1.00 1.00 426037 tit 5 36.53 0.18 16.15 0.08 41.79 0.21 0.99 0.99 426037 tit 6 36.78 0.27 16.18 0.12 41.76 0.31 1.00 1.00 426008 apa 1 27.93 0.06 15.36 0.04 40.96 0.09 0.99 0.99 426008 apa 2 37.36 0.04 15.95 0.02 39.13 0.05 0.97 0.96 426008 apa 3 38.28 0.03 15.91 0.01 38.70 0.04 0.93 0.85 426008 apa 4 38.68 0.08 16.08 0.03 38.73 0.08 0.99 0.98 426008 apa 5 64.84 0.94 21.41 0.31 35.92 0.52 1.00 1.00 wr = whole-rock, grt = garnet, ky = kyanite, tit = titanite, apa = apatite. r1 = 206pb/204pb vs.207pb/204pb error correlation (ludwig 1988). r2 = 206pb/204pb vs.208pb/204pb error correlation (ludwig 1988). * errors are two standard deviations absolute (ludwig 1988). geus bulletin 6.pmd 10-02-2005, 09:5470 71 (3) (2) wr (6) (7) (1) (1) 426050 garnet 426050 garnet 426050 kyanite 426050 kyanite a 826 + 96 ma (mswd = 0.14) 20 7 pb / 20 4 pb (5) 0 100 200 300 c 20 7 pb / 20 4 pb 20 8 pb / 20 4 pb (3) 2010 24 28 32 36 40 70 60 50 40 30 500 400 300 200 100 0 20 8 pb / 20 4 pb all datapoints: 876 + 93 ma (mswd = 11.2) (3) (3) (4) 426050 monazite 426050 monazite e 20 7 pb / 20 4 pb 206pb / 204pb 206pb / 204pb 0 100 200 300 500 400 300 200 100 0 20 8 pb / 20 4 pb 34 38 26 30 22 14 18 10 17.0 16.2 16.6 15.8 15.4 15.0 36 28 32 24 16 20 12 (3) (5) (7) (6)(2) (1) b 0 100 200 300 (4) (2) (5) (6) (1) d (3) (4) (4) (5) (6) ▲ ▲ ▲ ▲ ▲ ▲ ▲ (2) ▲ ▲ ▲▲▲ ▲ ▲ ▲ ▲ ▲ ▲ wr mon azi te tre nd garnet trend zircon trend (3) (4) (3) f 0 100 200 300 (4) wr ▲wr garnet kyanite (4) (4) 2010 24 28 32 36 40 901 + 42 ma (mswd = 1.86) fig. 2. uranogenic (207pb/204pb – 206pb /204pb) and thorogenic vs. uranogenic (208pb/204pb – 206pb /204pb) pb isotope diagrams with pbsl data from step-leaching experiments on garnet (a, b) and kyanite (c, d), from mica schist sample 426050 (krummedal supracrustal sequence). e and f are steps representing monazite inclusions within the garnet and kyanite. geus bulletin 6.pmd 10-02-2005, 09:5471 72 (fig. 2a) yielding a 207pb/206pb date of 876 ± 93 ma (mswd = 11.2). the 208pb/204pb vs. 206pb/204pb diagram (fig. 2b) reveals the presence of mineral inclusions in the garnet. the different th/u ratios of the host mineral and the inclusions may explain the large mswd value of the errorchron. steps 3 and 4 have very high th/u ratios, interpreted as representing monazite inclusions (th/u > 3; dewolf et al. 1996). all the monazite is leached out in step 4, causing the observed drop in the th/u ratio. the very low th/u ratio in steps 5 and 6 is characteristic of zircon leach steps (th/u < 1; dewolf et al. 1996). all the zircon is dissolved in step 6. step 7 was undertaken because of the red colour of the residue after step 6, showing that garnet was still present. the only leach steps dominated by garnet are the two first, where all the most primitive pb is extracted, and step 7. these three steps together with the wholerock analysis yield an isochron date of 826 ± 96 ma (mswd = 0.14). the large error of the date is due to the low precision of step 7. the same procedure was carried out on kyanite, and again there is evidence for the presence of both monazite and zircon inclusions (fig. 2d). steps 3 and 4 are dominated by monazite, and step 6 by zircon. three steps (1, 2 and 5) are interpreted as representing kyanite, but while the individual analyses are very precise they do not form a sufficiently wide spread in the pb ratios to yield a precise date. the pb wholerock analysis and the kyanite-dominated steps define a slope which yields an isochron date of 1219 ± 790 ma (mswd = 0.24). given the large uncertainty, this date does not yield any useful chronological information. the monazite-dominated steps from the garnet (3 and 4) and the kyanite (3 and 4) plot on a linear trend in both the 207pb/204pb vs. 206pb/204pb and 208pb/ 204pb vs. 206pb/204pb diagrams (fig. 2e, f). the four monazite steps yield an isochron date of 901 ± 42 ma (mswd = 1.86). the monazite and garnet dates are in general accordance with the ion microprobe analyses of metamorphic zircon rims from the krummedal supracrustal sequence that have yielded neoproterozoic ages around 940 ma (thrane et al. 1999a, b; kalsbeek et al. 2000). the cores of detrital zircons from the same study yielded ages ranging from c. 1100 to 1900 ma, and it therefore serves no practical purpose to calculate an age from the zircon steps, as these will represent a mixture of ages. sample 422766, also derived from the krummedal supracrustal sequence, was collected by j.c. escher and k.a. jones in the southern part of andrée land, very close to the contact with the structurally underlying crystalline basement (fig. 1). the sample contains garnet and kyanite crystals up to 5 cm in diameter. the kyanite was analysed by pbsl, while the garnet was considered too altered to justify analysis. seven steps were undertaken on the kyanite, and the analyses represent an almost perfect leaching pattern (fig. 3); all fall on a linear array in both the 207pb/204pb vs. 206pb/204pb and 208pb/204pb vs. 206pb/204pb diagrams, except for step 7 which has a lower 208pb/204pb ratio that probably indicates the presence of zircon inclusions. a 207pb/206pb date of 437 ± 62 ma (mswd = 2.6) is obtained using all the steps, while if step 7 is excluded a date of 445 ± 58 ma (mswd = 3.2) is obtained. the large error is due to the analytical error of steps 4 and 5. (3) (7) (2) (6) (1) (3) (7) (2) (6)(1) 422766 kyanite 28 24 20 16 12 a 445 + 58 ma (mswd = 3.2) 20 7 pb / 20 4 pb (5) (4) 0 40 80 120 160 200 240 422766 kyanite 400 300 200 100 0 b 20 8 pb / 20 4 pb (5) (4) 0 40 80 120 160 200 240 206pb / 204pb fig. 3. uranogenic (207pb/204pb – 206pb /204pb) and thorogenic vs. uranogenic (208pb/204pb – 206pb /204pb) pb isotope diagrams with pbsl data from step-leaching experiments on kyanite, from mica schist sample 422766 (krummedal supracrustal sequence). geus bulletin 6.pmd 10-02-2005, 09:5472 73 crystalline basement the east greenland caledonian orogen is dominated by major thrust sheets of reworked orthogneiss complexes. the crystalline basement is divided into an archaean terrain to the south of 72°50′n and a palaeoproterozoic terrain to the north (fig.1; thrane 2002). pbsl analyses on titanite from a metagabbroic gneiss (426018) in the archaean crystalline basement complex west of innermost forsblad fjord (fig. 1) were undertaken. this gabbroic gneiss has yielded a smnd model age (t dm ) of 3.25 ga (thrane 2002). the whole-rock analysis has the same pb ratios as step 1, indicating that the whole-rock and titanite are in equilibrium. the six leach steps together with the wholerock analysis yield an isochron date of 504 ± 48 ma (mswd = 1.81). however, the 208pb/204pb vs. 206pb/ 204pb diagram (fig. 4b) suggests that the titanite contains small amounts of monazite inclusions, which result in slightly elevated th/u ratios for steps 3 and 4 compared with the titanite trend. if steps 3 and 4 are excluded, an isochron date of 486 ± 15 ma (mswd = 0.16) is obtained (fig. 4a). titanite from a sheared garnet amphibolite (426037) cutting the basement gneisses of nathorst land (fig. 1) was also analysed. all six data points define an isochron date of 335 ± 140 ma (mswd = 0.11; fig. 4c); the large error of the date is due to the limited spread in the data points, as well as the large analytical errors of steps 3 and 4. in the 208pb/204pb vs. 206pb/ 204pb diagram (fig. 4d) the analyses show an unusual (3) (2) (wr) (6) (1) (1) 426018 titanite 426037 titanite a 486 + 15 ma (mswd = 0.16) all steps: 335 + 140 ma (mswd = 0.11) 20 7 pb / 20 4 pb (5) 0 40 80 120 160 200 c 20 7 pb / 20 4 pb 206pb / 204pb 206pb / 204pb 20 8 pb / 20 4 pb (3) 4010 50 70 90 50 46 42 33 34 70 60 50 40 30 20 8 pb / 20 4 pb (4)23 25 19 21 17 15 13 19 20 17 18 16 15 14 0 20 40 60 80 100 (3) (2) wr (6) (1) b (5) 0 40 80 120 160 200 (4) 426018 titanite (3) (2) (6) (1) d (5) (4) 426037 titanite (4) (5) (6) ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ (2) ▲ ▲ ▲ ▲ fig. 4. uranogenic (207pb/204pb – 206pb /204pb) and thorogenic vs. uranogenic (208pb/204pb – 206pb /204pb) pb isotope diagrams with pbsl data from step-leaching experiments on titanite from a gabbroic gneiss in the basement (a, b; sample 426018) and a garnet amphibolite (c, d; sample 426037). geus bulletin 6.pmd 10-02-2005, 09:5473 74 pattern: step 2 has an elevated th/u ratio compared to the general trend while step 3 has a lower th/u ratio. these features cannot be explained by the presence of monazite and zircon inclusions. if steps 2 and 3 are excluded from the isochron an even less precise date of 309 ± 230 ma (mswd = 0.03) is obtained. apatite from a tonalitic basement gneiss (426008) collected west of innermost forsblad fjord (fig. 1) was also analysed. zircon crystals from this sample have yielded u-pb ages of c. 2800 ma (thrane 2002). apatite dissolves much more easily than silicate phases, so a weaker acid and shorter leaching times were used in this experiment. the analyses form a complex pattern (fig. 5). step 1 is too thorogenic to derive from apatite, and it is interpreted instead as influenced by allanite since this is very easily dissolved and has a higher th/u ratio than apatite. step 2 is less thorogenic than step 1, but more so than step 3, and is therefore interpreted as a mixture between allanite and apatite. step 3 is the only step dominated by apatite. the only possible way to obtain a date is thus by combining the whole-rock analysis and step 3, which yields a date of 392 ± 24 ma (fig. 5a). step 1 falls on the isochron, while step 2 falls slightly above, demonstrating that the two mineral phases were almost in equilibrium, with the presumed allanite being slightly older corresponding to its higher closure temperature. in the 208pb/204pb vs. 206pb/204pb diagram it seems that both steps 3 and 4 are apatite steps, but in the 207pb/ 204pb vs. 206pb/204pb diagram it is clear that step 4 is older and must be influenced by zircon inclusions which were leached out in the strong acid of step 5. the whole-rock analysis and step 5 yield a date of 2159 ± 46 ma. summary and discussion the analyses reported in this paper are the first pbsl analyses reported on rocks from the caledonian orogen of east greenland. all the samples have been analysed only once. several of the dates obtained are not consistent with existing ages from the area, and some of the new dates are also somewhat controversial; replicate analyses should therefore be made for all the samples, to confirm that the dates are consistent, before any definite interpretations can be made. thus different interpretations are presented in the discussion that follows. the reliability of the pbsl method is still an open question. the main concern is the importance of the micro-inclusions contained in the mineral being analysed, and whether it is possible to be certain which combination of minerals is dissolved and affect the individual steps. this important point has not yet been resolved, and must be kept in mind when evaluating the new dates. supracrustal rocks the pbsl study demonstrates that neoproterozoic monazite and garnet are present in the krummedal supracrustal sequence; evidence of caledonian monazite has previously been reported (kalsbeek et al. 2000). zoned garnets have often been recorded (elvevold & gilotti 1999; thrane et al. 1999b), of which 23 23 19 17 15 13 a 426008 apatite 392 + 24 ma 20 7 pb / 20 4 pb 42 40 38 36 34 b 426008 apatite 20 8 pb / 20 4 pb 206pb / 204pb (wr) (wr) (1) (2) (5) (4) (3) (1) (2) (3) (4) (5) 0 20 40 60 70 0 20 40 60 70 fig. 5. uranogenic (207pb/204pb – 206pb /204pb) and thorogenic vs. uranogenic (208pb/204pb – 206pb /204pb) pb isotope diagrams with pbsl data from step-leaching experiments on apatite from a tonalitic basement gneiss (sample 426008). geus bulletin 6.pmd 10-02-2005, 09:5474 75 the outer rims are interpreted to be caledonian whereas there has previously been doubt as to whether the cores were neoproterozoic or early caledonian. in contrast, the presence of neoproterozoic kyanite has not been demonstrated in this study. petrographically it is often difficult to determine to which mineral paragenesis the kyanite belongs, and thus it cannot be ruled out that some of the kyanite in parts of the krummedal supracrustal sequence may be neoproterozoic (elvevold & spears 2000). evidence of early caledonian metamorphism in the crystalline basement? the closure temperature for titanite is estimated by dahl (1997, and references therein) to be in the range of 620–680°c, and by cherniak (1993) in the range of 575–707°c, depending on the grain size. the titanite date of 486 ± 15 ma for sample 426018, together with the date of the monazite inclusions, suggest that the crystalline basement did experience caledonian medium to high-grade metamorphism. no other ages of c. 486 ma have yet been obtained in east greenland. the age of the caledonian collision in east greenland is usually referred to the interval 430–425 ma, on the basis of zircon ages from granite intrusions and the time of migmatite formation in the krummedal supracrustal sequence (watt et al. 2000; kalsbeek et al. 2000, 2001). no comparable zircon ages have been recorded in the crystalline basement rocks in the study area, where evidence of the caledonian overprint is restricted to imprecise lower concordia intercept ages ranging from 467 ± 18 ma to 443 ± 25 (thrane et al. 1999a). it is not possible to determine whether these lower intercept ages correspond to the ‘traditional’ east greenland caledonian range of events, or to a potential earlier event. in north-east greenland caledonian zircons have been recorded in some palaeoproterozoic gneisses (kalsbeek et al. 1993), which is in line with the assumption that the crystalline basement complexes of this northern region were more strongly reworked during the caledonian orogeny. it might be speculated that the titanite date of 486 ± 15 ma is a cooling age, while the slightly older monazite micro-inclusions in the titanites could represent the peak of a collision event – comparable to the early caledonian event in scandinavia (mørk et al. 1988; andréasson 1994, 2000). however, this is not possible in east greenland, since ordovician carbonates were still being deposited in the iapetus-margin basin that lay east of the laurentian crystalline basement at this time; there is no associated clastic input that would be expected if a collision had taken place nearby. the exceptionally thick ordovician carbonate succession in east greenland (smith 1991) is indicative of a significant increase in the rate of subsidence, and it is possible that the c. 500 and 486 ma dates are instead related to extension. the c. 430 ma ages are thus still the best indication of the main caledonian collision phase in east greenland. apatite, yielding the youngest caledonian date of 392 ± 24, could be interpreted to represent the time where the basement gneisses cooled to < 500°c (dahl 1997). acknowledgements the isotope data described in this paper were acquired at the geological institute, university of copenhagen. robert frei is thanked for introducing and guiding me in the pb-pb step-leaching method. jan c. escher and kevin a. jones kindly provided sample 422766. critical comments on the manuscript by adam a. garde, a.k. higgins and feiko kalsbeek are greatly appreciated. this project was based on funding from the danish natural science research council. minik rosing and martin whitehouse are thanked for reviewing the manuscript. references andréasson, p.g. 1994: the baltoscandian margin in neoproterozoic – early palaeozoic times. some constraints on terrane derivation and accretion in the arctic scandinavian caledonides. tectonophysics 231, 1–32. andréasson, p.g. 2000: finnmarkian deep-seated imbrication of a margin of baltica: evidence from high-grade deformation zones in the kebnekaise mts., swedish caledonides. 24. nordiske geologiske vintermøte, trondheim, norway, 6–9 january, 2000. abstract volume. geonytt 1, 32 only. berger, m. & braun, i. 1997: pb-pb dating of apatite by a stepwise dissolution technique. chemical geology 142, 23–40. cherniak, d.j. 1993: lead diffusion in titanite and preliminary results on the effects of radiation damage on pb transport. chemical geology 110, 177–194. dahl, p.s. 1997: a crystal-chemical basis for pb retention and fission-track annealing systematics in u-bearing minerals, with implications for geochronology. earth and planetary science letters 150, 277–290. dewolf, c.p., zeissler, c.j., halliday, a.n., mezger, k. & essene, e.j. 1996: the role of inclusions in u-pb and sm-nd garnet geus bulletin 6.pmd 10-02-2005, 09:5475 76 geochronology: step-wise dissolution experiments and trace uranium mapping by fission track analysis. geochimica et cosmochimica acta 60, 121–134. elvevold, s. & gilotti, j.a. 1999: reaction history of metapelitic rocks from goodenough land, east greenland. in: frederiksen, k.s. & thrane, k. 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(eds): the caledonide orogen: scandinavia and related areas, 1115–1124. chichester: john wiley & sons. smith, m.p. 1991: early ordovician conodonts of east and north greenland. meddelelser om grønland geoscience 26, 81 pp. thrane, k. 2002: relationships between archaean and palaeoproterozoic basement complexes in the southern part of the east greenland caledonides: an ion microprobe study. precambrian research 113, 19–42. thrane, k. 2004: palaeoproterozoic age of the basement gneiss complex in the charcot land tectonic window, east greenland caledonides. in: higgins, a.k. & kalsbeek, f. (eds): east greenland caledonides: stratigraphy, structure and geochronology. geological survey of denmark and greenland bulletin 6, 57–66 (this volume). thrane, k., kalsbeek, f. & watt, g.r. 1999a: evidence for a grenville event in the east greenland caledonian fold belt. in: frederiksen, k.s. & thrane, k. (eds): second symposium on east greenland geology, mainly caledonian. abstract volume. danmarks og grønlands geologiske undersøgelse rapport 1999/21, 37 only. thrane, k., watt, g.r., kinny, p.d., jones, k.a. & escher, j.c. 1999b: early neoproterozoic breakup of rodinia: sims u-pb ages from the east greenland caledonides. eug 10, terra abstracts 4, 119 only. todt, w., cliff, r.a., hanser, a. & hofmann, a.w. 1993: recalibration of nbs lead standards using a 202pb + 205pb double spike. terra abstracts 5 (supplement 1), 396 only. watt, g.r., kinny, p.d. & friderichsen, j.d. 2000: u-pb geochronology of neoproterozoic and caledonian tectonothermal events in the east greenland caledonides. journal of the geological society (london) 157, 1031–1048. geus bulletin 6.pmd 10-02-2005, 09:5476 geological survey of denmark and greenland bulletin 23, 2011, 37-40 37 the brackish baltic sea and the more saline kattegat are connected by three straits, lillebælt, storebælt and øresund (fig. 1). of the three straits, lillebælt is the narrowest, with 700 m at its narrowest point, widening out towards the south to around 25 km (fig. 2). in the narrow parts of lillebælt, water depths around 30–50 m are common. in the northern part of lillebælt the depth is 16–18 m and in the southern part the depth is around 35 m. storebælt and øresund have played important roles as outlets during the history of the baltic sea, and their histories have been much discussed (björck 1995; bennike et al. 2004). in contrast, lillebælt has received little attention. in this paper we present 11 new radiocarbon accelerator mass spectrometry (ams) ages and propose a curve for holocene relative shore-level changes in lillebælt. we use the term shore-level changes rather than sea-level changes because we have constructed both lake-level and sea-level changes. during the last deglaciation of the lillebælt region, large channels were eroded by northward-flowing subglacial meltwater. these channels are now found at the bottom of the strait, and most of them are kept free of sediments by strong bottom currents. however, late and postglacial sediments are found in some parts of the channels. several submerged settlements have been reported from the lillebælt region (andersen 1985). they are dated to the mid-holocene from artefacts and by radiocarbon dating. methods combined high-resolution, sub-bottom profiling and sediment coring were carried out from r/v alexander von humboldt. the seismo-acoustic equipment included a sediment echosounder (fig. 3), and the profiles obtained were used for the selection of the core sites (fig. 4). a 6 m long vibrocorer was used for coring. we also had access to vibrocores and seismic profiles from a survey conducted in connection with a planned gas pipeline. this material was handed over to the geological survey of denmark and greenland (geus) from dansk olie og naturgas a/s. selected cores with the most complete stratigraphy were sub-sampled for studies of macrofossils. the samples were wet sieved and analysed using a dissecting microscope. repostglacial, relative shore-level changes in lillebælt, denmark ole bennike and jørn bo jensen 10°e 14°e 56°n 57°n sweden denmark germany kattegat storemecklenburg bælt femern bælt baltic sea bucht kieler lillebælt øresund fig. 2 bucht 50 km jylland fyn fig. 1. map of denmark and surrounding area showing the location of lillebælt and the other straits connecting the baltic sea to kattegat, as well as place names mentioned in the text. fig. 2. bathymetry of lillebælt with the location of the vibrocores indicated. © geus, 2011. geological survey of denmark and greenland bulletin 23, 37–40. open access: www.geus.dk/publications/bull 55°30´n 10°e 10°e mr009a la010 281900 la006b 281910 la003 fyn jylland als fig. 3 55°30´n 55°n 0–10 10–20 20–30 >30 water depth (m) 10 km 3838 mains of plants and marine molluscs were submitted for ams radiocarbon dating (table 1). several published dates were also used for the reconstruction of shore-level changes (table 2; k-samples are conventional ages, and the lus-sample is an ams age). we have used a reservoir age of 400 years for the marine samples, however, the reservoir age may have varied somewhat during the holocene (olsen et al. 2009). sediments, palaeoecology and chronology the oldest sediments consist of till that shows an internal, chaotic reflection pattern and a sharp upper boundary. a few cores also penetrated meltwater sand. till and meltwater sand accumulated during the last glaciation and deglaciation of the region. the glacial deposits are locally overlain by late glacial sediments, which are found in the channels. the late glacial sediments show conformable internal reflectors, and consist of clay, silt and fine-grained sand. one sample has been dated to 11 400–11 900 cal. years bp, corresponding to a late younger dryas age (table 1, poz-8924). we suggest that the late glacial sediments were partly deposited in a branch of the baltic ice lake. in the deeper parts of lillebælt, black, organic-rich sediments are widespread. the sediments are commonly laminated and may contain abundant fragments of small roots and fruits of telmatic plants. some of this sediment is swamp peat, but most of it is coarse detritus gyttja. the organic-rich sediments are usually overlain by laminated calcareous gyttja clay. however, in core la006b lake sediments are found below peat. this succession is interpreted as overgrowing of a basin. samples from the lake deposits gave ages of c. 11 000–8800 cal. years bp (early holocene, table 1, poz-5754, poz-5755, poz-5753, poz-8859, poz8860). 27 22 17 d ep th ( m b .s .l. ) marine deposits early holocene freshwater deposits late glacial deposits glacial deposits la003 la003 281910 281910 281900 281900 la006b la006b 1 km a b fig. 3. a: original seismic profile, obtained by a sediment echosounder. b: interpretation below. for location see fig. 2. cores labelled 2819xx were collected from r/v alexander von humboldt, and cores labelled la were collected for dansk olie og naturgas a/s. 11000 7100 11600 10600 281900 c la y si lt sa n d c la y si lt sa n d c la y si lt sa n d c la y si lt sa n d h ii h i gl h i lg ? ? 7600 281910 ss ss ss ss 11000 la003 core 10800 la006b l it h o lo gy st ru ct u re s d ep th ( m b el o w s ea l ev el ) 25 26 27 28 29 24 ss radiocarbon age lamination structureless irregular lamination bioturbation silt/clay/gyttja/marl sand pebbles peat heterolithic fig. 4. sedimentological logs from vibrocores from the lillebælt. radiocarbon ages are in calibrated calendar years bp. gl: glacial. lg: late glacial. h i: early holocene freshwater. h ii: brackish and marine holocene deposits. 39 marine sediments from protected areas consist of laminated or bioturbated, fine-grained, organic-rich mud. sandy and silty sediments are found in shallow water areas and in areas with strong bottom currents (fig. 3). shells and shell fragments of marine molluscs are common. on the acoustic records, the marine deposits are mostly transparent or show continuous reflectors parallel to the lower boundary. in three cores we dated the lowermost shell of marine molluscs we could find. the oldest age determination is c. 7700 cal. years bp (table 1, poz-5790). in core la010 bioturbated sand is present in the upper part of the core. the fauna implies brackish conditions. a sample from the bottom of the sand unit was dated to c. 8500 cal. years bp (table 1, poz-5767). we suggest that the sand marks the first marine influence in the area. shore-level changes on the basis of the available radiocarbon ages, we have reconstructed relative shore-level changes in the region (fig. 5). the relative shore level was low during the early part of the holocene and probably rose slowly throughout the early holocene, and at the same time a large lake existed in the area. as the shore level rose this lake increased in size and at around 8500 cal. years bp it was transformed into a brackish water body. two dates from core mr009a provide an important fix point for the shore-level evolution (table 1). the dates show that a peat now found 9 m below sea level was transgressed by the sea between c. 8200 and c. 7700 cal. years bp, and around 8000 cal. years bp marine conditions were established. later sea-level changes are constrained by six published radiocarbon dates (table 2). they comprise two dates of wood from marine gyttja, two dates from ostrea edulis shells, an age from a bone found in a grave at a water depth of 2.7 m and a bone of harp seal from a submarine settlement. the two latter dates come from sites that were situated above the contemporary sea level. discussion in lillebælt, late glacial sediments are found in incised channels. the younger dryas sequence that consists of finegrained laminated clay and silt is followed by a hiatus which was probably formed during the final drainage of the baltic ice lake, when shore level dropped around 25 m over a few years (björck 1995). the maximum shore level of the baltic ice lake in the south-western baltic sea was around 20 m b.s.l., and this lake may have extended as far west as southwestern kieler bucht (jensen et al. 2002). the baltic ice lake may also have extended into southern lillebælt. core laboratory species* sediment depth age ( c calibrated no. no. b.s.l. (m) years bp) age (years bp)§ la003 poz-5754 m. trifoliata, c. mariscus detritus gyttja 28.10–28.20 9670 ± 50 10 789–11 210 p. australis la006b poz-5755 m. trifoliata, c. mariscus lake marl 26.02–26.03 9460 ± 50 10 567–11 068 la010 poz-5767 b. albae, c. mariscus brackish sand 16.60–16.70 7700 ± 70 8384–8599 la010 poz-5753 m. trifoliata, c. mariscus lake gyttja 16.72–16.78 7880 ± 50 8556–8976 mr009a poz-5790 mytilus edulis marine mud 9.70 7280 ± 40 7842–7649 mr009a poz-5805 b. albae peat 9.80–9.90 7420 ± 50 8074–8372 281900 poz-8820 arctica islandica marine sand 25.00–25.05 6590 ± 40 6994–7225 281900 poz-8859 p. tremula, b. nana lake clay 26.50–26.60 9350 ± 50 10 419–10 702 281900 poz-8860 p. tremula, b. albae detritus gyttja 26.80–26.88 9670 ± 50 10 789–11 210 281900 poz-8924 salix sp. clay 26.98–27.08 10 110 ± 60 11 401–11 910 281910 poz-8821 m. edulis, m. balthica marine sand 25.25–25.26 7140 ± 40 7528–7692 * full names are: menyanthes trifoliata, cladium mariscus, phragmites australis, betula sect. albae, betula nana, populus tremula, mytilus edulis, maco ma balthica. §calibration is according to the intcal09 dataset (terrestrial samples) and the marine09 dataset (marine samples). table 1. new radiocarbon ams age determinations from lillebælt 14 terrestrial peat lake deposit brackish sediment marine fossils marine gyttja bone from grave seal bone from settlement d ep th ( m b el o w s ea l ev el ) 12 10 8 6 4 2 0 mr009a age (cal. ka bp) mr009a 281900281910 la006b la003 m m sg b b m g s 0 10 20 30 la010 281900 fig. 5. curve showing relative shore-level changes in southern lillebælt during the holocene. green: lake phase. red: brackish water phase. blue: marine phase. ka: 1000 years. 4040 during the earliest holocene, large parts of lillebælt were dry land, but local bogs and lakes must have existed in the deeper parts. as the shore level began to rise, local lakes and bogs became widespread. during continued shore-level rise, bogs were transformed into lakes, and a large lake developed in the southern part of lillebælt. it was connected to another large lake to the south in kieler bucht, and to other large lakes in femer bælt, mecklenburg bucht and storebælt. later, the ongoing eustatic sea-level rise led to brackish and then to marine conditions in lillebælt. the first marine influence was via storebælt when southern lillebælt was a fjord. however, the fjord was transformed into the lillebælt strait during continued rapid sea-level rise. the oldest dated marine shell from lillebælt is from 7700 cal. years bp, but brackish water conditions are suggested at 8600–8384 cal. years bp. the youngest lake deposits (around 17 m below sea level) are dated to 8976–8556 cal. years bp. in storebælt, the oldest dated marine shell gave an age of 8100 cal. years bp (bennike et al. 2004), and in the mecklenburg bucht, the oldest shell date is c. 8000 cal. years bp (rößler et al. 2011). the early holocene deposits in lillebælt show no indication of a lowering of the shore level before being inundated by marine waters. conclusions glacial till and holocene marine deposits are widespread in lillebælt. in the deeply incised channels late glacial and early holocene non-marine deposits are found, these units are separated by an erosional boundary. the late glacial deposits were probably deposited during pre-allerød and allerød times, as well as during the younger dryas. the early holocene non-marine deposits have yielded ages between 11 000 and 8800 cal. years bp. the late glacial unit consists of lake deposits, and we suggest that the baltic ice lake extended into southern lillebælt. during the early holocene, a large lake existed in southern lillebælt; this lake expanded in size during shorelevel rise. the oldest shell of a marine mollusc from lillebælt is dated to 7700 cal. years bp, but brackish conditions were probably established at around 8500 cal. years bp. acknowledgement the captain and crew of r/v alexander von humboldt, and in particular the cruise leader, the late wolfram lemke are thanked for their help during the marine cruise. references andersen, s.h. 1985: tybrind vig, a preliminary report on a submerged ertebølle settlement on the west coast of fyn. journal of danish archaeology 4, 52–69. bennike, o., jensen, j.b., lemke, w., kuijpers, a. & lomholt, s. 2004: lateand postglacial history of the great belt, denmark. boreas 33, 18–33. bennike, o., rasmussen, p. & aaris-sørensen, k. 2008: the harp seal (phoca groenlandica erxleben) in denmark, southern scandinavia, during the holocene. boreas 37, 263–272. björck, s. 1995: a review of the history of the baltic sea, 13.0–8.0 ka bp. quaternary international 27, 19–40. jensen, j.b., kuijpers, a., bennike, o., laier, t. & werner, f. 2002: new geological aspects for freshwater seepage and formation in eckernförde bay, western baltic. continental shelf research 22, 2159–2173. olsen, j., rasmussen, p. & heinemeier, j. 2009: holocene temporal and spatial variation in the radiocarbon reservoir age of three danish fjords. boreas 38, 458–470. petersen, k.s. & rasmussen, k.l. 1996: the impact of radiocarbon datings on natural historical sciences in denmark: especially paleozoological and shore-line datings. pact 49, 117–130. rößler, d., moros, m. & lemke, w. 2011: the littorina transgression in the southwestern baltic sea: new insights based on proxy methods and radiocarbon dating of sediment cores. boreas 40, 231–241. doi: 10.1111/j.1502-3885.2010.00180.x. laboratory material depth age ( c calibrated reference no. b.s.l. (m) years bp) age (years bp)* k-3558 human bone 2.7 6740 ± 80 7459–7727 andersen (1985) k-4150 alnus wood 4.7 6380 ± 100 7153–7480 andersen (1985) k-4149 tilia wood 2.8 5370 ± 100 5922–6317 andersen (1985) k-5680 ostrea edulis shells 5.0 5940 ± 70 6645–6995 petersen & rasmussen (1995) k-5681 ostrea edulis shells 4.0 5780 ± 70 6445–6797 petersen & rasmussen (1995) lus-6136 phoca groenlandica bone 2.0 5595 ± 50 5885–6144 bennike et al. (2008) *calibration is according to the intcal09 dataset (terrestrial samples) and the marine09 dataset (marine samples). table 2. published radiocarbon age determinations from lillebælt 14 authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: obe@geus.dk geological survey of denmark and greenland bulletin 31, 2014, 47-50 47 stormwater management: methods for measuring near-surface infiltration capacity in clayey till britta bockhorn, marina bergen jensen and knud erik s. klint glacial till forms a major proportion of the surface deposits in northern europe, and in denmark more than 40% of the land surface is covered by clayey till. at the same time the majority of densely populated areas are situated on this fertile sediment type. in urban areas, one of the major tools in adaptation to climate change are sustainable drainage systems (suds). their function is to manage the increasing amounts of stormwater on site, often by direct infiltration into the sediment. accordingly, a realistic estimate of nearsurface hydraulic properties is required when dimensioning suds for infiltration. clayey tills are generally believed to have a low-bulk hydraulic conductivity and thus a low infiltration capacity. however, clayey tills can be very heterogeneous and especially their bulk hydraulic properties can vary significantly depending on the distribution of permeable structures such as macropores (e.g. earthworm holes and fractures) and sand lenses within the till matrix (klint & gravesen 1999; nilsson et al. 2001; kessler et al. 2012). the saturated hydraulic conductivity (ksat) of clayey tills varies from c. 1.0 × 10–10 to c. 1.0 × 10–4 m/s and thus covers a significant span (fredericia 1990; mckay et al. 1993). assessment of this value is a major challenge when considering the variability of this sediment type. it is therefore important to determine how representative standard infiltration tests are, specifically in tills, when used to formulate infiltration strategies. goals and scope in this study we compare three different methods for measuring ksat close to the surface: the double ring infiltrometer, the guelph permeameter and infiltration tests in a small excavation. each of these methods represents different scales and depths using different flow mechanisms. the goal of the study is to: 1. evaluate the suitability of these methods to return realistic ksat values in tills, taking into account the geological heterogeneity of a clayey till from infiltration scale (0.5 m × 0.5 m) to field scale (100 m × 100 m), and 2. suggest relevant scales and strategies for infiltration tests in future experiments. © 2013 geus. geological survey of denmark and greenland bulletin 28, 47–50. open access: www.geus.dk/publications/bull fig. 1. a: map of denmark showing the location of the test site. b: map of the test site showing the distribution of various sediment types and the locations of the infiltration tests. a b10°e 57° 57°n 55° 50 km 20 m 55°39.5´n 12°16.6´e postglacial clay clayey till sandy clayey till sandy till postglacial sand fill deeper borehole auger test guelph permeameter measurement double ring infiltrometer measurement shallow (2.5 and 4 m) monitoring borehole infiltration hole høje tåstrup 4848 field experiments the infiltration tests were conducted on former agricultural land next to the technological institute in høje taastrup, denmark (fig. 1). the area represents a site with a typical danish clayey till. the geological setting is dominated by two basal till beds overlying glaciofluvial deposits and flow tills deposited in a supraglacial environment. the glacial deposits overlie highly fractured limestone bedrock c. 14–16 m below the ground surface. the primary groundwater table is located in the limestone bedrock at depths greater than 16 m and a secondary groundwater table is found in the upper till unit. the latter is located around 3 m below the surface during summer and around 1.0 to 0.5 m below the surface during winter. an area of c. 100 m × 100 m was mapped in great detail using a hand auger (jakobsen et al. 2011) with sampling in a 10 × 10 m grid. shallow boreholes were cored to depths of 2.5 –4 m and used for monitoring the annual fluctuation of the secondary groundwater table. two deeper boreholes were cored to a depth of c. 16 m. two large holes were excavated to 5 and 8 m below the ground surface and used for detailed mapping of fractures and collection of large, intact samples for hydraulic tests in the laboratory. the guelph permeameter method measures the steadystate rate of water flow required to maintain a constant depth of water in a 40 cm deep and 8 cm wide cylindrical borehole. water flows out of the outlet tube through a perforated section located above the permeameter tip. the guelph reynolds & elrick (1985) 60 cm infiltration holedouble ring infiltrometerguelph permeameter direction of measured water flow expenditure of time (time required to achieve full saturation) 3-dimensional 1-dimensional 3-dimensional effect on soil structure smearing during excavation causing potential underestimation of k. counteraction: careful removal of smearing with knife non-invasive highly invasive suitability to represent geological heterogeneity 1–2 hours depending on soil moisture 1–2 hours depending on soil moisture poor good degreeof disturbance 1.5 hours to several hours depending on soil moisture 8 cm 5–10 cm 30.5 cm 60.5 cm minimally invasive rather poor counteraction: conduct at least a triplet of measurements counteraction: conduct as many measurements as possible 100 cm 200 cm literature on estimation of ksat reynolds et al. (2002) 40 cm formation of cracks during insertion of cylinders into the soil creation of preferential flow routes causing potential overestimation of k. counteraction: sealing of contact with clay on the outside of the inner cylinder smearing during excavation of hole causing potential underestimation of k. counteraction: careful removal of smearing with knife or use of excavation technologies that prevent or minimise smearing, e.g. chain excavator fig. 2. overview and principle of the different techniques used in this study. the blue arrows represent water flow. k: hydraulic conductivity. 49 permeameter method is based on the assumption of threedimensional steady-state infiltration from a cylindrical test hole into the sediment. two concentric metal cylinders with diameters of 30.5 and 60.5 cm were used for the double ring infiltrometer method. after removal of the sward, the cylinders were carefully pressed 5–10 cm into the sediment. water was poured into the inner cylinder, and also into the outer cylinder to prevent lateral movement of water beneath the inner cylinder, thus maintaining one-dimensional flow conditions. the amount of cumulative infiltration with time under fallinghead conditions was recorded and ksat values determined. the infiltration holes were excavated to a depth of 60 cm with an inner area of 100 × 200 cm. smearing caused by the excavation process was carefully removed with a knife. the holes were filled with water and when a steady state was attained, the infiltration rate from the hole into the sediment was measured directly. a total of 41 infiltration tests were conducted across the site: 19 guelph permeameter measurements, 18 double ring infiltrometer measurements and four infiltration tests in the excavated holes. more information on the methods is provided in fig. 2. results and discussion the application of conventional infiltration technologies indicates that the saturated hydraulic conductivity (ksat) of tills is a spatially highly variable property. in two guelph permeameter measurements no infiltration at all was observed, which might be due to smearing during preparation of the borehole or compaction by heavy machines as the field site is former agricultural land. compaction might also be the reason for one no-flow measurement in the double ring infiltrometer. the results are presented in table 1 and fig. 3. the guelph permeameter and the double ring infiltrometer average values are lower than those from the infiltration holes. this is probably due to not fully saturated conditions around the holes as the area is rather large compared to the area used for the guelph permeameter and double ring infiltrometer measurements, where saturated conditions are attained reasonably quickly. the fact that the holes involve a much larger area than the guelph permeameter and the double ring infiltrometer means that they involve more than one sediment type and accounts best for macro-pores, confirming that geological heterogeneity on clayey till plains influences the hydraulic conductivity even at infiltration plot scales. a realistic estimate of hydraulic conductivity is crucial when planning how to manage stormwater infiltration. the present field-based study shows that hydraulic conductivity values are strongly influenced by the physical scale of the field measurements. this is mainly due to the local distribution of macro-pores and how well the geological heterogeneity is represented. the importance of scale-dependent variability of ksat has also been demonstrated in previous studies (e.g. jenssen 1990; ronayne et al. 2012), but is still widely neglected when hydraulic conductivity values are used for stormwater management practices. due to the large variation of grain-size distribution in tills, it is not recommended to apply methods that define ksat based on grain-size distribution. neither is it recommended 1 × 10–9 1 × 10–8 1 × 10–7 1 × 10–6 1 × 10–5 guelph permeameter double ring infiltrometer infiltration holes sa tu ra te d hy dr au lic c on du ct ivi ty (m /s) fig. 3. box and whisker diagram illustration of the range of saturated hydraulic conductivity values obtained using the guelph permeameter, double ring infiltrometer and infiltration hole techniques. the scale is logarithmic. whiskers are plotted at the maximum and minimum values, however, the minimum value for the double ring infiltrometer test is below the scale and not shown. guelph permeameter 19 9.12 × 10–8 6.18 × 10–6 1.44 × 10–6 double ring infiltrometer 18 7.43 × 10–13 9.7 × 10–6 8.26 × 10–7 infiltration holes 4 1.4 × 10–6 1.46 × 10–5 7.25 × 10–6 table 1. summary statistics of saturated hydraulic conductivity values method number mininum maximum arithmetic mean 5050 to use ksat values obtained with a double ring infiltrometer or a guelph permeameter alone as they can vary by several orders of magnitude already on a infiltration plot scale. in our study we found a variable of more than two orders of magnitude. data from infiltration holes give more realistic values. however, they are highly invasive and it may be difficult to excavate adequate holes in densely populated areas. instead of using highly invasive infiltration holes, we recommend to carry out combined hydrogeological investigations where double ring infiltrometer and guelph permeameter measurements are supported by geological information from maps of near-surface deposits and borehole descriptions. in that less-destructive way, small-scale geological heterogeneity can be revealed and the most suitable areas for stormwater infiltration can be selected to enhance work efficiency of infiltration devices. acknowledgement the work was conducted as part of the innovation consortium cities in waterbalance (byer i vandbalance) financed by the danish council for technology and innovation. references fredericia, j. 1990: saturated hydraulic conductivity of clayey tills and the role of fractures. nordic hydrology 21, 119–132. jakobsen, p.r., hermansen, b. & tougaard, l. 2011: danmarks digitale jordartskort. danmarks og grønlands geologiske undersøgelse rapport 2011/40, 28 pp. jenssen, p.d. 1990: methods for measuring the saturated hydraulic conductivity of tills. nordic hydrology 21, 95–106. kessler, t.c., klint, k.e.s., nilsson, b. & bjerg, p.l. 2012: characterization of sand lenses embedded in tills. quaternary science reviews 53, 55–71. klint, k.e.s. & gravesen, p. 1999: fractures and biopores in weichselian clayey till aquitards at flakkebjerg, denmark. nordic hydrology 30, 267–284. mckay, l.d., cherry, j.a. & gillham, r.w. 1993: field experiments in a fractured clay till. 1. hydraulic conductivity and fracture aperture. water resources research 29, 1149–1162. nilsson, b., sidle, r.c., klint, k.e., bøggild, c.e. & broholm, k. 2001: mass transport and scale-dependent hydraulic tests in a heterogeneous glacial till–sandy aquifer system. journal of hydrology 243, 162–179. reynolds, w.d. & elrick, d.e. 1985: in situ measurement of field-hydraulic conductivity, sorptivity, and the alpha-parameter using the guelph permeameter. soil science 140, 292–302. reynolds, w.d., elrick, d.e., youngs, e.g., amoozegar, a., booltink, h.w.g. & bouma, j. 2002: saturated and field-saturated water flow parameters. in: dane, j.h. & topp, g.c. (eds): methods of soil analysis, part 4, physical methods, 797–878. madison, wi: soil science society of america. ronayne, m.j., houghton, t.b. & stednick, j.d. 2012: field characterization of hydraulic conductivity in a heterogeneous alpine glacial till. journal of hydrology 458–459, 103–109. authors’ addresses b.b. & m.b.j., department of geosciences and natural resource management, university of copenhagen, rolighedsvej 23, dk-1958 frederiksberg, denmark. e-mail: bboc@ign.dk.ku k.e.s.k., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. geological survey of denmark and greenland bulletin 23, 2011, 21–24 21 lowand intermediate-level radioactive waste from the danish nuclear research facility, risø, includes construction materials from the reactors, different types of contaminated material from the research projects and radioactive waste from hospitals, industry and research institutes. this material must be stored in a permanent disposal site in denmark for at least 300 years (indenrigsog sundhedsministeriet 2007). the ministry of health and prevention presented the background and a decision plan for the danish parliament in january 2009 (ministry of health and prevention 2009) and all political parties agreed to the plan. in the beginning of 2011 three studies were presented to the parliament (http://www.im.dk/aktuelt/nyheder/ forebyggelse/2011/maj/slutdepot.aspx): (1) a pre-feasibility study for the final disposal of radioactive waste, geological characterisation of potential disposal areas for radioactive waste from risø, denmark peter gravesen, merete binderup, bertel nilsson and stig a. schack pedersen 22: ålbæk 15: randers 11: hindsgavl– fænø 13: thyrsted–glud 12: klejs–sønderby 10: kertinge mark 9: vemmenæs 8: mid7: langeland 4: rødbyhavn 3: falster south 6: stevns 2: hammeren– vang 5: risø 14: lysnet bakker16: limfjord 21: skive 17: hvidbjerg 18: harre vig 19: branden–junget 20: thise lol land falster møn 1: østermarie– paradisbakkerne west south bornholm s jæl land fyn jy l land denmark germany sweden langeland south fjord 10°e 14°e 55°n 57°n 56°n 50 km fig. 1. map of denmark showing the location of the 22 selected areas. red: the six best areas. blue: the 16 remaining areas. © geus, 2011. geological survey of denmark and greenland bulletin 23, 21–24. open access: www.geus.dk/publications/bull 2222 (2) a study on radiation doses from the transport of radioactive waste to a future repository and (3) a study on identifying potential disposal areas. the latter study was conducted by the geological survey of denmark and greenland (geus) and the aim was to locate a sediment or rock body with low permeability down to 100–300 m below the ground surface. the ultimate goal is longterm protection of people and environment by isolating the radioactive waste in a final depository. this goal can be reached by identifying a significant volume of sediments or rocks characterised by a low flow regime and high absorption potential. geus was given the task to locate approximately 20 potential disposal areas. geological setting and data requirements in denmark, many types of fine-grained sediments and crystalline rocks occur from the ground surface down to a depth of 300 m. descriptions of these sediments and rocks are based on existing information and include four main types: (1) granite and gneiss on bornholm, (2) chalk and limestone, (3) fine-grained palaeogene and neogene clay and (4) quaternary clayey till and clay. in europe, the most studied geological formations for disposal of radioactive waste are clay (belgium, france, germany and switzerland), crystalline rocks (sweden, finland and switzerland) and salt (germany). salt diapirs and salt pillows as well as deep-seated basement rocks are not included in the current study. several types of existing data were compiled for the preliminary selection of the approximately 20 potential areas as outlined by indenrigsog sundhedsministeriet (2007). the recommendations follow the guidelines of the international atomic energy agency (iaea 1994, 1999, 2005). the study provides an overview of the distribution of various deposits and tectonic features in denmark. tectonic features, the distribution of layers of low permeability and the distribution of fractured sediments and rocks are important for the assessment and selection of areas suitable for disposal of radioactive waste (gravesen et al. 2010). data collection and compilation the deposits are described from the ground surface down to a depth of at least 100 m. the description of each area comprises: (1) geological conditions such as general geology, surface geology and profiles, sediment and rock characteristics, tectonic features and structures, seismic activity, geological and structural models, ground stability; (2) hydrogeological conditions such as groundwater characteristics, vulnerable drinking-water supply areas, geoand hydrochemical conditions; (3) ground-surface conditions: terrain and topography, surface processes, climate and climate changes, restrictions and limitations in connection with protection of nature areas; (4) summary of area conditions with final remarks and literature. the area descriptions address the following important issues (iaea 1994, 1999, 2005): (1) the final disposal site should be situated in an area with homogeneous geological conditions. it should be demonstrated that these conditions are found with a high degree of probability at the selected site. as the geological conditions in many parts of denmark are heterogeneous on both local and regional scales, the goal was to find a sufficiently large area with continuous and homogeneous sediment or rock bodies without fractures or preferential flow paths. (2) the geological deposits shall contribute to isolate the radioactive waste. this is most effective if the disposal site is underlain or surrounded by low permeability layers such as clay, silt or crystalline rock. (3) restriction of pore water flow from the disposal site is favoured by deposits of low permeability. (4) the disposal site must be placed at the greatest possible distance from the nearest groundwater aquifer. (5) the disposal site must be located outside areas of special drinking-water interests. (6) the surrounding groundwater aquifer must be able to contribute to dilution of any radioactive material that might leak from the disposed material. (7) the surrounding sediments, rocks and ground water aquifer material must have a high potential for absorption of any leaking radioactive components. (8) geological processes at the terrain surface should not have any influence on the quality of the disposal site. criteria and methods for selection the potential disposal areas fit the criteria set up by gravesen et al. (2010, 2011) to various degrees. due to the heterogeneous geological conditions in denmark it is not possible to fulfil all the criteria within one area. therefore, it is necessary to assess the criteria and to compare the areas. criteria – the geological and hydrogeological criteria have been central for the selection of the areas. the type of disposal site had not been decided when the study started and the following criteria were considered the most important: (1) the deposits from the ground surface and downwards should be as homogeneous and of as low permeability as possible. this means that highly permeable deposits such as sand and gravel should be of limited extent. (2) the deposits of low permeability that enclose the waste should comprise thick layers of large horizontal extent. 23 other criteria were also important for the selection: (1) areas of special drinking-water interests have been totally avoided. areas of drinking water interests comprise a large part of denmark and it was impossible to totally exclude these areas. (2) we tried to avoid groundwater bodies of good status. (3) eu nature protection areas (natura 2000) were completely avoided; other nature and heritage protection areas were avoided if possible. (4) large cities and suburbs of larger cities were also avoided. methods for selection – the methods used were as follows: (1) the geological conditions in denmark have been evaluated based on existing data and information (e.g. pedersen 1989, håkansson & pedersen 1992). it was decided which types of sediments and rocks should be included in the work. the sediments and rocks fulfilling most of the demands are: crystalline rocks (granite and gneiss), some types of limestone, fine-grained pre-quaternary clay and fine-grained quaternary clay. if possible, layers of clayey tills should cover these deposits. (2) the next step was to identify and avoid the areas of special drinking-water interests, natura 2000 areas and large cities; these areas were not treated further. (3) the remaining areas were analysed according to the criteria of thick deposits of low permeability with large horizontal extent from the ground surface and downwards. possible future climate changes were also considered during the selection process. the current prognoses from the intergovernmental panel on climate change cover c. 100 years. the waste facility must exist for at least 300 years; predictions of climate changes for the latter part of this period are highly uncertain. the sea level is expected to rise 0.5–1.0 m, but it may rise more. survey results our survey resulted in the selection of 22 areas throughout denmark (fig. 1). the compilation and scientific evaluation of the data resulted in several new findings concerning the danish geology. data from a large number of boreholes have been compiled from areas that normally attract little interest from geologists and water-resource managers. in addition, the selected areas are of no or little interest for local drinkingwater exploitation, nature protection or archaeology. the selected areas are larger than required for the final depository. therefore the final site can be located according to other limitations or parameters within the selected area. from 22 to six areas – it is suggested that the work of the next phase be concentrated on six areas, chosen from among the 22 areas. the 22 areas were chosen so that they fulfil the criteria or most of the criteria. they are all qualified areas, according to an evaluation based on existing data and knowledge, but the amount of information and knowledge varies from area to area. the six areas are assessed as being slightly better than the remaining 16 areas, which are reserve areas that can be included if the six areas cannot be used for some reason h ig h 3: falster south 7: langeland south 9: vemmenæs 8: mid-langeland 11: hindsgavl 12: klejs–sønderby 13: thyrsted–glud 14: lysnet bakker 15: randers fjord g la ci o te ct o n ic i n flu en ce 16: vorde–lyndelse (ls) 18: harre vig (?) 1: paradisbakkerne 5: risø (incl. meltwater clay) 4: rødbyhavn 11: fænø 10: kertinge mark 16: handbjerg (ls) 17: hvidbjerg 19: branden–junget 20: thise 21: skive west sm al l 22: ålbæk thickness of glacial clay <30 m >30 m 3: falster south > 50 m 7: langeland south 8: mid-langeland 13: thyrsted–glud 11: hindsgavl 14: lysnet bakker t h ic kn e ss o f p al ae o ge n e cl ay 12: klejs 15: randers fjord 18: harre vig 19: branden–junget < 50 m many structures in palaeogene clay few 21: skive west 20: thise 17: hvidbjerg 10: kertinge mark 4: rødbyhavn 5: risø (including fractures in chalk) 9: vemmenæs 11: fænø 12: sønderby fig. 2. diagram showing the thickness of glacial clay (mainly clayey till) versus the importance of glaciotectonic influence. the green box show the best localities and the red localities are the selected ones. ls: limfjord south. fig. 3. diagram showing the structures in palaeogene clay versus the thickness of palaeogene clay. the green box show the best localities and the red localities are the selected ones with palaeogene clay, the sixth locality is in crystalline rock (østermarie–paradisbakkerne). the localities hammerenvang and stevns are not shown in the diagrams, because they have no palaeogene clay or they have not been influenced by glaciotectonics. 2424 that ‘overrides’ the geological criteria. the overriding factors could be grounded in infrastructure and regional planning as well as protests from citizens. results from future detailed field and laboratory studies may also lead to the rejection of some areas, which were considered suitable according to existing knowledge. the six areas are østermarie–paradisbakkerne in bornholm’s regionskommune, rødbyhavn in lolland kommune, kertinge mark in kerteminde kommune, hvidbjerg in struer kommune, thise in skive kommune and skive west also in skive kommune (fig. 1). the geological criteria used to select the six potential areas were: (1) the areas have clay or crystalline rocks of low permeability from the ground surface to at least 100 m depth with only few subsurface structures. (2) clayey till dominates the upper part, but this covering layer of clayey till is relatively thin (less than 30 m) in most of the area, and layers of low permeability are rapidly reached. this meets the requirements for a medium deep repository (30–100 m), where the depository should be surrounded by layers of low permeability. (3) glaciotectonic influence has little significance (down to 30 m) but this is often difficult to assess for a large area, because glaciotectonic features can only be demonstrated from outcrop or borehole data. (4) in the selected areas, only areas of no, limited or some drinking-water interests are found. (5) in the areas, there are no significant occurrences of groundwater bodies with good status. (6) the terrain is mostly flat and of little relief, and the landscape is considered stable and without risk of landslides. the different properties of the 22 areas were compared. areas with glacial till less than 30 m thick and palaeogene clay over 50 m thick were considered the best areas (figs 2, 3). also areas with only minor glaciotectonic influence and areas with only few tectonic structures in the palaeogene clay were ranked as best. areas with glaciotectonic deformations such as fractures, faults and folds (klint & gravesen 1999) are often characterised by sand and gravel occurring between finer grained layers. with respect to tectonic structures in the palaeogene clay, it should be noted that some areas are poorly covered by data. as far as possible, the selected areas are located in regions with limited or no drinking-water interests. however, some of the selected areas are located partly or wholly in areas of drinkingwater interests. six of the areas are located in areas with quaternary clay, basement rock or chalk/limestone and do not include palaeogene clay. these areas have relatively thin top layers of clayey till (less than 20 m). fractures and tectonic structures may occur. based on our analysis of the properties we conclude that the six named areas are better suited for radioactive waste disposal than the other 16 areas. final remarks our survey resulted in the selection of 22 areas throughout denmark. six of these areas are preferred on geological and hydrogeological criteria. eventually, the six areas will be reduced to one, two or three areas that appear promising in which further detailed field work will be carried out. the field investigations include analysis of the geological, hydrogeological, hydrochemical and geomechanical conditions. finally, one site will be chosen for the final waste disposal. acknowledgement financial support was provided by the parliament of denmark. references gravesen, p., nilsson, b., pedersen, s.a.s. & binderup, m. 2010: lowand intermediate radioactive waste from risø, denmark. location studies for potential disposal areas. report no. 1. data, maps, models and methods used for selection of potential areas. danmarks og grønlands geologiske undersøgelse rapport 2010/122, 47 pp. gravesen, p., nilsson, b., pedersen, s.a.s. & binderup 2011: lowand intermediate radioactive waste from risø, denmark. location studies for potential disposal areas. report no. 11. områdebeskrivelser – description of areas. dansk og engelsk resume. danmarks og grønlands geologiske undersøgelse rapport 2011/51, 64 pp. håkansson, e. & pedersen, s.a.s. 1992: geologisk kort over den danske undergrund. 1:500 000. copenhagen: varv. (map sheet). iaea 1994: siting of near surface disposal facilities. safety guides. safety series 111-g-3.1, 37 pp. iaea 1999: near surface disposal of radioactive waste. requirements. safety standards series ws-r-1, 29 pp. iaea 2005: borehole facilities for the disposal of radioactive waste. specific safety guide. safety standards series ssg-1, 102 pp. indenrigsog sundhedsministeriet 2007: beslutningsgrundlag for et dansk slutdepot for lavog mellemaktivt affald, 47 pp. unpublished report, indenrigsog sundhedsministriet, copenhagen, denmark. klint, k.e.s. & gravesen, p. 1999: fractures and biopores in weichselian clayey till aquitards at flakkebjerg, denmark. nordic hydrology 30, 267–284. ministry of health and prevention 2009: redegørele om beslutningsgrundlag for et dansk slutdepot for lavog mellemaktivt affald, 13 pp. report, ministry of health and prevention, copenhagen, denmark. pedersen, s.a.s. (ed.) 1989: jordartskort over danmark 1:200 000. four map sheets: nordjylland; midtjylland; sydjylland; fyn, sjælland, øer og bornholm. copenhagen: danmarks geologiske undersøgelse. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: pg@geus.dk hydrocarbon gases in palaeogene volcanic rocks from the lopra-1/1a well, faroe islands 79© geus, 2006. geological survey of denmark and greenland bulletin 9, 79–90. available at: www.geus.dk/publications/bull hydrocarbon gases in palaeogene volcanic rocks from the lopra-1/1a well, faroe islands troels laier hydrocarbon gases were monitored in the drilling fluid during deepening of the lopra-1 well from 2178–3565 m, in which thermogenic, methane-rich gases had been found previously. the mud gas concentration, up to 105 ppm of methane, was generally higher in the hyaloclastite sequence, 2470 m – terminal depth (td), than in the overlying lavas of the lower basalt formation. the highest concentrations of mud gas in the lower basalt formation were associated with the more porous tuffaceous zones, whereas no simple relationship could be established between measured mud gas concentrations and porosity of the hyaloclastic rocks, which showed less marked porosity variations than the lavas. chemical (c2+ < 1%) and isotopic (δ13c1: –34 to –39‰) compositions of seven samples of mud gas collected at peak gas concentrations between 2657 m and 3442 m compare well with those of the hydrocarbon gases which had been seeping more or less continuously into the existing well since 1983, suggesting a common origin of the gases. headspace methane concentrations measured in 135 canned samples of cuttings were scattered between 10 ppm and 6 × 103 ppm, with the exception of six samples from a short interval, 2685– 2745 m, which showed consistently high values > 104 ppm. no particularly gas-rich zones were indicated, however, by the mud gas, nor was any significant change in lithology noted for this interval. it is possible that the technique of turbo-drilling, that had been attempted over a short interval, 2657– 2675 m prior to collection of the high-level methane samples, may have caused enhanced degassing due to the very fine cuttings produced. chemical and isotopic composition of headspace gas and mud gas indicated the same type of gas throughout the well, although headspace methane tended to be more enriched with respect to the 13c isotope. the origin of the lopra-1 gas is discussed in the light of recent information obtained from source rock studies of central east greenland and the faroe–shetland basin. keywords: faroes, lopra-1/1a, volcanics hydrocarbon gas, isotopes, headspace methane, cuttings _______________________________________________________________________________________________________ geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: tl@geus.dk as hydrocarbons were not expected to occur in the basalts of the faroe islands, no monitoring of hydrocarbon gases in the drilling fluid was performed while drilling the 2178 m deep scientific lopra-1 well in 1981. a few observations made while drilling, however, indicated negligible amounts of gases in the basalts (waagstein et al. 1984). none of these observations suggested larger accumulations of gas so, after having reached td in november 1981, the drilling fluid in the well was replaced by fresh water. soon afterwards, the well began to flow approximately 9.5 litres/min. (k. højgård, personal communication 1982) and it was decided after a few days to shut-in the well. the lopra-1 well was not re-opened until march 1983, when temperature logging was to be performed. by that time the wellhead pressure had increased to 19.5 bars (balling et al. 1984) and an inflammable gas had accumulated within the 190 m long casing at the top of the well (p.h. nielsen, personal communication 1983). the volume of the gas was estimated to be roughly 9 m3 at 1 bar pressure (p.h. nielsen, personal communication 1983) geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1979 80 and gas chromatographic analysis of a sample collected at the wellhead showed the gas to consist of methane (72%) and nitrogen (27%) plus traces of higher hydrocarbons (jacobsen & laier 1984). after temperature logging, the well was shut-in again and an increase in wellhead pressure was noted shortly after. two months later, the wellhead pressure had increased to 10.8 bars (fig. 1). collection of a new gas sample showed that the composition of the gas was almost identical to that of the sample collected two months earlier. furthermore, isotopic analyses indicated that the gas was thermogenic in origin, the δ13c of methane being –39.6‰ (jacobsen & laier 1984). encouraged by the hydrocarbon discoveries of foinaven in 1992 and schiehallion in 1993 in the british sector 160 km south-east of the faroe islands, new investigations of the hydrocarbon traces in the lopra-1 well were carried out in 1992–1993, including stable isotopic analysis of gases and biomarker analyses of the small quantities of higher hydrocarbons extracted from the water flowing out of the well (laier et al. 1997). an additional gas sample was taken for isotopic analysis in july 1994, before the deepening of the lopra-1 well commenced in july 1996 (table 1). this paper presents the results of new chemical and isotopic analyses of both mud gas samples and headspace gas of canned cuttings collected during a deepening of the lopra-1/1a well in 1996. the results of the continuous mud gas readings are compared with well logs in order to identify any particular gas-rich zones, and the possible origin of the hydrocarbon gases in the basalts is discussed. sampling and methods mud gas gas samples were taken at the mud-logging unit on different occasions (table 2). the gas samples were stored in 100 cm3 steel bottles which had been flushed with mud gas for two minutes before closing the valves. the samples were sent to the geological survey of denmark and greenland (geus) within a week for chemical and isotopic analyses. the concentrations of hydrocarbon gases in the gas samples were analysed using a shimadzu gc9 gas chromatograph equipped with a flame ionisation detector (fid). separation of the gas constituents was performed on a 4 m 1/8” diameter ss column packed with silica gel 70/80 mesh using helium as carrier gas (20 ml/min.) using a temperature programme beginning at 100°c, increasing 20°c/min. to 250°c final temperature. headspace gas of canned cuttings unwashed cuttings were sampled at approximately 10 m intervals and stored in 0.5 litre cans to which bactericide had been added. the cans were sent to geus for headspace gas and isotopic analyses after the completion of the well. the cans were generally in a good condition when received by geus, except for a few cans where the lid was not sealed tightly. the cans were placed upside down, punctured through a septum and 0.2 ml of headspace gas fig. 1. wellhead pressure of the lopra-1 well during the shut-in period after temperature logging (p.h. nielsen, personal communication 1983). year ch4 c2h6 c3h8 ic4h10 nc4h10 n2 o2+ar co2 h2 δ13c1 δ13c2 δ13c3 δd1 1994 60.2 0.31 0.048 0.014 0.012 39.1 0.06 0.01 –32.1 –26.5 –150 1992 40.6 0.14 0.025 0.008 0.007 59.1 n.d. 0.01 –32.4 –26.5 –148 1983 1.9 0.41 0.064 0.015 0.018 27.5 0.04 0.01 –32.5 n.a. –133 n.a.: not analysed; n.d.: not determined. concentrations are given in vol%. stable isotopic ratios are given (in parts per thousand) relative to the pee dee belemnite (pdb) and standard mean ocean water (smow) standards. table 1. chemical and isotopic composition of gases from the original lopra-1 well collected at the wellhead <0.01 <0.01 <0.01 –39.9 –41.5 –39.6 0 4 8 12 date pr es su re ( kg /c m 2 ) 83-03-20 83-04-19 83-05-20 geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1980 81 results and discussion gases from the original lopra-1 well no casing had been set in the old lopra-1 well except for the upper 190 m of the borehole. therefore, the substantial amount of gas which had accumulated at the wellhead during the 16 months shut-in period after drilling, could in theory have entered from any part of the open hole. however, jacobsen & laier (1984) observed an increase in the gas to water ratio with depth suggesting that the gas entered mainly in the deeper parts of the well. the flow of gas (0.8 litre/min., ambient pressure and temperature) from the well measured by jacobsen & laier (1984) during their field work in june 1983 also suggested that the gas continued to seep into the open borehole of the lopra-1 well. taking into account the methane dissolved in the water flowing out of the well (at a water flow rate of 12.8 litres/min.), the total flow of methane is estimated to have been approximately 0.9 litre/min.; 0.56 litre/min. in the gas phase (72% ch4) plus 0.3 litre/min. dissolved in water. so the total flow of methane in june 1983 was over 1 m3 per day. if the influx of methane had been the same during the initial 16 months shut-in period, a much larger quantity of gas would be expected than the 9 m3 (ambient pressure) noted when the well was re-opened in march 1983. however, the flux may have been slower during the shut-in period due to higher pressure (19.5 bars) at the wellhead. gas was still seeping into the borehole 13 years after drilling of the lopra-1 well. the composition of the hydrocarbon gas had not changed significantly during that period (table 1) and this suggests a major single source of the gas. the source provided mainly methane, as the nitrogen content in the gas-to-water ratio measured at wellwas withdrawn using a gas-tight syringe to be analysed by gas chromatography as described above. if a high concentration of methane was found, 60 ml of headspace gas was transferred to an evacuated serum bottle for later isotopic analysis. the volume of headspace gas in the cans, generally between 100–150 ml, was checked by addition of water. no corrections on headspace gas concentrations were attempted to account for the difference in headspace volumes. isotopic analyses headspace gas and mud gas samples were transferred to a ¼” column packed with porepak q immersed in liquid nitrogen. separation of the gas constituents was then performed allowing the column to heat to ambient temperature. methane and ethane were combusted over cuo at 900°c and the resulting carbon dioxide was isolated in flame-sealed glass ampoules using cryogenic traps. isotopic measurements were performed on a finnigan mat 251 mass spectrometer at the university of copenhagen. the isotopic ratio is expressed in the usual delta notation relative to the pdb standard. δ13c(‰ ) = (rsample/(rstandard – 1)) × 1000 where r = 13c/12c the analytical precision of the isotopic measurements is ± 0.2‰ for larger quantities of gas (> 20 µl co2 stp), typical of the amounts of methane in all of the analysed samples, and ± 0.5‰ for smaller quantities of gas (< 10 µl co2 stp), typical of the ethane. interval (m) event comments 2360–2450 sour smell and low ph of drilling fluid 2657–2675 turbo drilling finer cuttings. possible enhanced degassing. potential risk of gas generation due to high temperature. 2680–2822 one shale shaker specific gravity increase, 1.11–1.18 g/cm3, of drilling fluid due to solids build-up. out of order 2992 20% dilution of mud reduction of specific gravity, 1.18–1.13 g/cm3. 3023 mud centrifuge installed reduction of specific gravity to 1.06–1.08 g/cm3. 3158 string got stuck 5 m3 diesel pill added. 3091–3565 side track, lopra-1a 1% oil in drilling fluid according to daily analysis. table 2. drilling events which may influence hydrocarbon gas measurements low ph may cause corrosion and generation of artificial gas. indication of bacterial activity. bacteria may generate/consume methane. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1981 82 np hi -0 .1 2200 2300 2400 2500 2600 2700 2800 2900 3000 3100 10 102 105 103 104 3100 3200 3300 3400 3500 3.2 2.6 2.0 rock density (g/cc) mud gas ch4 (ppm) 10 102 105 103 104 3.2 2.6 2.0 rock density (g/cc) mud gas ch4 (ppm) h ya lo cl as tit es , m os tly la pi lli -t uf fs (s ub aq ue ou s) lo w er b as al t fo rm at io n (s ub ae ri al ) fig. 2. rock density and mud gas recorded in the extended lopra-1/1a well. head can be accounted for by the amount of atmospheric nitrogen initially dissolved in meteoric water (at 6°c) percolating into the basalts. mud gas from the lopra-1/1a deepening the original lopra-1 well, drilled in 1981, was deepened in 1996. a 75/8” casing was first set in the existing 8½” borehole before the well was extended downwards using a 6½” drill bit. the lopra-1 well was terminated at 3158 m due to technical problems and a sidetrack was drilled from 3095 m to a td of 3565 m (lopra-1a). any hydrocarbons detected in the drilling fluid or in the sealed samples of drill cuttings of the extended well section must have come from the new drilling activities. no casing was set in the new well sections, which means that gases detected in the drilling fluid (fig. 2) represent the sum of gases released by drilling in addition to the gases that may have geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1982 83 seeped through the sidewall of the entire open hole. a kcl-polymer mud was used as drilling fluid and the well logs indicate that no substantial mud cake had been built up in the hole. therefore, gas seepage from the sidewall of the open hole was probably more significant for the lopra-1/1a well compared to conventional exploration wells in sedimentary rocks. the mud gas and headspace gas data from the lopra1/1a well are the only information available on hydrocarbon gases in the basalts penetrated, as no further tests were performed in the well after drilling. drilling was optimised to detect hydrocarbon gases in the rocks penetrated, i.e. the specific gravity of the drilling fluid was kept low and the trap for continuous hydrocarbon measurement was placed in an almost closed mudline system. however, the mud gas and headspace gas must be interpreted with caution. it is important to consider any possible effects of the drilling process itself. the drilling events which may have influenced the hydrocarbon measurements and which could give a false impression of the hydrocarbon variation in the rocks, have been listed in table 2, together with their possible implications. the specific gravity of the mud could not be kept constant until a centrifuge had been installed (when drilling had reached 3023 m), after which the specific gravity remained at 1.06–1.08 g/cm3. during drilling of the lower basalt formation from 2184 to 2470 m the specific gravity remained almost constant at 1.03–1.05 g/cm3. at greater depths, the cuttings became finer and the specific gravity of the mud was increased gradually to 1.18 g/cm3 at 2992 m by increasing the content of solids. an increase in mud weight suppresses degassing from the rocks, which lowers the gas concentration in the mud. on the other hand, an elevated solids content may increase gas concentration, depending on how fast the cuttings degas. with these uncertainties in mind, a more detailed interpretation of the mud gas will be presented below. mud gas concentration varied from less than 10 ppm to 90 000 ppm of methane down to 3120 m depth, showing a general increase from 2400 m to 3120 m (fig. 2). higher hydrocarbons, mostly ethane, constituted less than one percent of the total gas throughout the drilling of the well and will not be dealt with in detail in the following. the intervals with very low mud gas concentrations, < 10 ppm, were observed only in the lower basalt formation above 2470 m. comparing mud gas concentration versus depth with various well logs as well as changes in various drilling parameters (rate of penetration, mud density etc.) it appears that rock porosity and rock density are the most important rock properties affecting the mud gas concentration in the upper 300 m of the extended well (fig. 2). variations in rock density and porosity correlate closely, so only the most complete log, the density log, is shown on fig. 2. assuming that the gas is indigenous to the rock and not an artefact of drilling, the highest gas concentrations should be found in the more porous, less dense rocks, as is also indicated by the data for the upper 300 m of the extended lopra-1/1a well. however, this simple relationship is not observed for the deeper parts of the well. from 2610 m to 3080 m the density of the rock generally increases and porosity decreases, yet a general increase in mud gas concentration is observed in this same interval, an increase that may be explained by the increased sidewall surface area. this suggests that degassing through the sidewall of the hole contributes more to the mud gas compared to the gas released by the drilling process itself. the presumed gas seepage through the sidewall of the open hole makes it difficult to decide which part of the well actually contributes to the gases recorded in the mud. however, for the deepest part of the well, below 3430 m, changes in the mud gas concentration appear to be correlated positively with porosity (and inversely with rock density), which may indicate that gas is present even in the deepest rocks penetrated by the lopra-1/1a well. thus, it may be concluded that hydrocarbons are most likely present in all of the rocks penetrated, particularly in the more porous rocks. however, no particular hydrocarbon-rich zones and no cap rock were identified. samples of mud gas were collected at maximum gas concentrations as far as possible (table 3). three of the samples were taken when mud circulation was resumed after changing the drill bit, and the rest of the samples were taken at maximum concentration while drilling. the table 3. samples of mud gas collected from the lopra-1/1a well on various occasions event resumed mud circulation after changing of drill bit resumed mud circulation after changing of drill bit gas collected after short drilling break gas collected at peak concentration while drilling gas collected at peak concentration while drilling gas collected at peak concentration while drilling gas collected after having drilled 6 m after trip 2657 2822 2946 2989 3107.5 3343 3442 date 96/08/28 96/09/07 96/09/09 96/09/10 96/09/19 96/10/23 96/10/26 sample no. 1 2 3 4 5 6 7 depth (m) geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1983 84 concentrations of hydrocarbon gases in the mud gas samples were usually lower compared to mud gas concentrations recorded in the mud by the time of sampling (table 4) although the chemical composition of the gas was the same. the reason for this is not clear, but insufficient flushing of the 100 cc steel bottle during sampling could be one reason. the chemical and isotopic composition of the mud gas (table 4) is not very different from that of the wellhead gas collected from the old lopra-1 well (table 1) suggesting a common origin for the gases. headspace gas of sealed cuttings methane concentration in the headspace gas of the 135 canned samples of cuttings from the lopra-1/1a well varied from 7 ppm to over 33 000 ppm, the highest concentrations being measured at 2273 m and in the interval from 2685 to 2745 m (fig. 3). the sample at 2273 m was collected just after drilling through a 2–3 m thick layer of lapilli tuff with a distinctly higher porosity compared to the massive lavas above and below. a higher mud gas concentration was also noted for this porous layer (fig. 2). the methane concentration in the samples from the 2685– 2745 m interval is more than one order of magnitude higher than that of most other samples from the well. this interval corresponds to the upper part of a section of dominantly lapilli tuff (2610–2880 m) and could indicate that these rocks contain more gas than those at other levels. however, the mud gas concentration did not shift to higher values as one might have expected if the gas content was generally much higher in the rocks of this particular interval (fig. 3). furthermore, the well logs show no correlation between the increase in headspace methane and any change in rock properties, such as porosity or density. a shift from rotary drilling to turbo-drilling using a diamond bit took place at 2657 m, and turbo-drilling continued to 2675 m where rotary drilling was resumed. cuttings produced by turbo-drilling are generally much smaller than cuttings produced by normal rotary drilling. it is possible that the decrease in cutting size may have led to enhanced degassing and thereby higher headspace methane concentrations. furthermore, one of the two shale shakers was out of order due to motor failure during drilling from 2680 to 2822 m. this may be one of the reasons for the observed build up of solids in the drilling fluid, which, combined with the finer cuttings produced by turbo-drilling, may have been the cause of the significantly higher methane concentrations in the headspace gas. alternatively, the much higher concentration of methane measured in the headspace gas shortly after turbodrilling could be explained by generation of artificial gas due to the very high temperatures which often occur using this drilling technique. if, however, the gases were generated by some artificial process, an increase in mud gas methane concentration should be expected, which is not the case (fig. 3). furthermore, the stable isotopic ratio of the methane is not markedly different in the zone with high methane concentrations compared to headspace methane from other levels (fig. 3). ethene, with concentrations up to 39 ppm, was observed in a number of headspace samples, particularly from the upper 300 m of the extended well where the ph of the drilling fluid was relatively low (7–8 compared with a typical value of 10) according to the daily drill reports. unsaturates like ethene are very uncommon among light hydrocarbons in natural gas, so this constituent was most likely formed artificially, either during drilling (faber et facing page: fig. 3. summary of mud gas and headspace results of the extended lopra-1 well (a) and lopra-1a sidetrack (b). larger circles represent c1/c2 ratio corrected for artificial gas indicated by ethene (c2=). numbers are stable isotopic values of methane. table 4. chemical and isotopic composition of mud gas samples depth (m) 2657 2822 2946 2989 3107.5 3343 3442 n.d.: not detected. n.a.: not analysed. ch4 ppm c2h6 ppm 31.6 31.7 19.0 81.6 74.4 17.3 28.4 c3h8 ppm 4.06 3.02 1.56 6.78 9.34 3.44 5.86 ic4h10 ppm 0.44 n.d. n.d. 0.36 0.92 0.66 0.98 nc4h10 ppm 0.63 n.d. n.d. 0.40 1.36 1.10 1.45 δ13c1 (‰) –38.6 –38.2 –36.7 –37.3 –38.1 –34.2 –36.0 δ13c2 (‰) –31.8 –35.0 –32.3 –33.4 –33.6 c1/c2 140 152 126 128 137 158 151 n.a. n.a. 4420 4820 2400 10440 10200 2740 4290 geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1984 85 a mud gas (ppm) 10 102 103 104 105 10 102 103 104 105 headspace gas (ppm) ch4 ch4 2200 2400 2600 2800 3000 3200 3000 3200 3400 3600 –38.2 –36.7 –33.7 –30.0 –33.8 –35.1 –28.6 –34.0 mud gas (ppm) 10 102 103 104 105 10 102 103 104 105 headspace gas ch4 ch4 0.1 1 10 102 headspace gas c1/c2 c2=(unsat.) δ13c –35.8 –35.0 –35.1 –39.2 –34.2 –36.0 d ep th ( m ) d ep th ( m ) δ13c –36.6 0.1 1 10 102 headspace gas c1/c2 c2=(unsat.) b 1 10 102 103 1 10 102 103 –38.6 –37.3 –38.1 –36.6 –36.7 geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1985 86 al. 1988) or as a result of corrosion (laier 1999). samples of cuttings generally contain iron filings from the drilling. unsaturates can be generated by a catalytic reaction of the fischer-tropsch type, and may occur in canned cuttings with relatively low ph, as was demonstrated for the swedish deep gas well at lake siljan (laier 1999). the concentration of ethene decreased when ph had been raised to over 10, by adding caustic soda and bactericide to the drilling fluid below 2500 m (fig. 3). the methane-to-ethane ratio of the gas may be useful in determining the origin of the gas in the volcanics, but a correction should be made for the artificial gas which may influence this ratio. the methane-to-ethane ratio of artificial gas varied between 3–5 in the siljan well (laier 1999), which is much lower than that of the gases in the lopra1/1a well (fig. 3). equal amounts of ethene and ethane were noted in the artificial gas of the siljan well (laier 1999); assuming a similar ratio of 1:1 for the artificial gas component in the lopra headspace samples, a correction has been made (fig. 3). headspace samples with methane concentrations below 100 ppm have generally lower methane-to-ethane ratios, but this cannot be taken as evidence of a different origin of the gas. this is more likely due to uncertainties of the headspace method as well as the analytical uncertainty for the very low ethane concentrations. the methane-to-ethane ratio varied between 102 and 103 for headspace gas samples with higher methane concentrations (> 100 ppm) that have less analytical uncertainty (fig. 3). the methane-to-ethane ratio remained fairly constant, around 102, in samples from the side track lopra1a. thus, given the uncertainties of the headspace method, it may be concluded that the hydrocarbon gas of the sealed cuttings mainly has the same origin as that of the mud gas and the previous wellhead gas. origin of the gasses the chemical and isotopic data for the gasses from the original lopra-1 well and the extended well sections have been plotted in the well-known classification diagram for hydrocarbon gases (fig. 4). the mud gas of the extended well plots fairly close to the wellhead gases of the old lopra-1 well, suggesting that they have a common origin. headspace gases, however, exhibit a much larger variation and are generally slightly more enriched with respect to carbon-13. such enrichment could result from either bacterial oxidation of methane (coleman et al. 1981) during storage or fractionation due to diffusion. the most carbon-13 enriched gases are found among the headspace samples with the highest methane concentration (fig. 3), therefore bacterial oxidation is unlikely to have been responsible for the carbon-13 enrichment. if degassing from the cuttings occurred mainly via diffusion, the gas may have fractionated due to degassing prior to storage in the tight cans. the hydrocarbon gases are thermogenic in origin, and the relatively low c2+ content indicates either a gas-prone source rock or a high maturity of the source rock. comparing the stable isotope values (δ13c) of methane through propane (tables 1, 4) with the δ13c versus source maturity lines published by rooney et al. (1995), a highly mature marine source rock (type ii kerogen) is most likely for the lopra gases. this is more obvious for the wellhead gases than for the mud gases as can be seen from fig. 5. the larger variation in δ13c among the mud gas samples compared to the wellhead gases, particularly for ethane, may be due to the much smaller quantities of hydrocarbons making the mud gases more susceptible to sampling and analytical errors. rooney et al. (1995) estimated that the reservoir gases used to construct the δ13c versus maturity line for the type (ii) kerogen had been generated in the temperature range 170–190°c, the upper part of which is not very different from the maximum palaeotemperature estimated for the base of the lopra-1/1a well (glassley 2006, this volume). furthermore, the reservoir gases used by rooney et al. (1995), which had isotopic values in the same range as those of the lopra-1 gases, were also very dry (m.a. rooney, personal communication 2000). on the other hand, using δ13c versus source maturity lines constructed by other researchers such as faber (1987) and berner & faber (1996), a somewhat lower maturity is 105 104 103 102 10 1 –80 –60 –40 –20 0 2 3 bacterial thermal type ii typ e i ii artificial wellhead δ13cch4 c 1/ (c 2 + c 3) mud gas fig. 4. plot of c1/c2 + c3 vs. δ13cch4 for lopra-1/1a gases. small solid circles represent headspace gas. (diagram modified from faber et al. 1999.) geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1986 87 predicted for the source of the lopra-1 gases (fig. 5). therefore, the maturity of the source rock that generated the gases is still not certain. the presence of a source rock is, however, very likely as was indicated by the continued and fairly significant flux of gas into the original lopra-1 well over a 13-year period. the gases entering the open lopra-1 well most likely came from the more porous layers of tuffs beneath the massive lavas of the lower basalt formation. no layers of coal were penetrated by the extended lopra-1/1a well, so the source is most likely located below the hyaloclastites that were extruded subaqueously (waagstein 2006, this volume). traces of higher hydrocarbons the traces of higher hydrocarbons measured in mud samples derived mainly from various additives (bojesen-koefoed & nytoft 2006, this volume) and cannot give any clues as to the possible origin of the lopra gases. fluid inclusion studies by konnerup-madsen (2006, this volume) do, however, indicate that hydrocarbon fluids were present in the volcanics at one time. but analysis of a fluid extracted from one calcite sample containing fluid inclufig. 5. cross-plot of isotope values (δ13c) of methane vs. ethane of wellhead and mud gas samples from the lopra-1/1a well. maturity lines, from faber (1987) and rooney et al. (1995), were based on reservoir data, whereas those of berner & faber (1996) were obtained from laboratory experiments. numbers attached to wellhead samples refer to sampling year. palaeogene basalt/ nansen fjord formation and faroes lower formation cretaceous–palaeogene sediments precambrian basement east greenland inland ice kangerlussuaq nansen fjord b l o s s e v i l l e k y s t scoresby sund faroe islands 100 km lopra 36° 33° 30° 66° 67° 68° 69° 70° 36° 33° 30° 27° 70° 69° 68° 67° 24°27° fig. 6. pre-drift reconstruction of central east greenland and the faroes block, modified from l.m. larsen et al. (1999). the arrow indicates the location of organic-rich sediments having ro = 1.2% (m. larsen, personal communication 2000). mud gas 83 94 iii = –3 0‰ (b ern er & fa ber 19 96 ) –40 ii = –2 5‰ (bern er & fa be r 1 99 6) –30 –20 ii = –30‰ –50 –40 –30 –20 iii (r oo ne y et a l. 1 99 5) ii (r oo ne y et a l. 1 99 5) iii = –2 5‰ 92 iii wellhead δ13cethane (‰) δ13 c m et ha ne ( ‰ ) geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1987 88 sion could not be related unambiguously to a particular source (bojesen-koefoed & nytoft 2006, this volume). waxes associated with zeolites reported from other parts of the faroe islands (laier et al. 1997) appeared to have been generated by a less mature source than that of the lopra-1/1a gases. thus, the presence of traces of higher hydrocarbons supports the assumption that hydrocarbons migrate or migrated in the volcanics of the area, but no direct connection between these higher hydrocarbons and the lopra gases have been documented. source rocks of the faroe islands area the scientific lopra-1 drilling was termined in november 1981 at 2178 m for technical reasons without having fulfilled its primary objective of penetrating the substratum of the lower basalt formation. the deepening of the lopra-1/1a well in 1996 to 3565 m penetrated hyaloclastites, mostly lapilli tuffs, under the base of the lower basalt formation at about 2470 m, but gave no definite clue as to the type of rocks occurring beneath the volcanics. the fact that the volcanics below the lower basalt formation were deposited subaqueously suggests that sediments were present in the area prior to the volcanic eruptions. such sediments would probably be similar in age to the sediments exposed beneath the basalts in the kangerlussuaq area in east greenland reported by m. larsen et al. (1999), sediments that probably give us the best clue as to what type of source rocks might exist beneath the volcanics at lopra, since the palaeodistance between the faroes block and central east greenland was only 100–120 km before the break-up of the northern north atlantic (fig. 6). the close connection between the two areas is emphasised by the similarity in chemical composition of the oldest basalts on either side of the northern north atlantic (l.m. larsen et al. 1999). the organic-rich rocks of the kangerlussuaq area range in age from late aptian to late paleocene. they proved to be post-mature with respect to hydrocarbon generation fig. 7. oil and gas fields west of shetland. far oe-s he tla nd c han ne l far oe-s he tla nd c han ne l 6°00’w 8°00’w 4°00’w 2°00’w 62°00’n 61°00’n 100 km foinaven shetland islands orkney islands clair schiehallion far oe– she tla nd c han ne l faroe islands lopra-1/1a geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1988 89 at most localities examined by m. larsen et al. (1999). however, a vitrinite reflectance value of ro = 1.2 measured on a late paleocene lacustrine mudstone (toc = 7%) 22 km north of nansen fjord (fig. 6) shows that source rocks with hydrocarbon generation potential do exist beneath the basalts in this area. an estimate of organic matter maturation related to burial depth alone may be obtained from the nansen fjord area which was covered by approximately 6 km of volcanic rocks prior to uplift (l.m. larsen et al. 1999). the total thickness of the faroe islands volcanics may exceed the 6.5 km presently known from exposures and drillings (waagstein 2006, this volume); however, a more precise estimate of the thickness is difficult to give. it seems realistic to assume that possible source rocks below the faroe islands volcanic succession still have some potential for hydrocarbon generation, given the close connection between the two areas prior to continental split-up. thus, the more or less continuous seepage of methane-rich gas into the open lopra-1 well during 1983–1994 may originate from a highly mature source rock located beneath the volcanics in the area. hydrocarbon migration from the faroe–shetland basin hydrocarbons may also have migrated into the faroe islands area from the faroe–shetland basin, though little is known of the possible migration pathways. the foinaven and schiehallion oil fields, 160 km south-east of the lopra-1 well, are the nearest known hydrocarbon occurrences around the faroe islands (fig. 7), but hydrocarbons are likely to have been generated closer to the faroe islands in basalt-covered rift basins beneath the faroe shelf (fig. 8). subsequently, these hydrocarbons may have migrated towards the faroe islands area via sandy turbidites located under the volcanics or via intrabasaltic sandstones. redeposition of coarse-grained sediments onto the faroe shelf during mid-paleocene uplift of east greenland has been inferred by m. larsen et al. (1999) but such sediments are not indicated on the geological profile of fig. 7, which shows only strata recognised on seismic sections. since much of the faroe–shetland basin is highly mature thermally, its lack of large gas accumulations has been explained by the presence of very oil-prone source rocks that have only little potential for later gas generation (scotchman et al. 1998). this assumption was based on kinetic studies of middle and upper jurassic kerogens from a number of exploration wells west of shetland. if that is the case, the assumption that the gases observed in the lopra-1 well came from the faroe–shetland basin is not supported by observations from this basin obtained within a reasonable distance from the faroe islands. conclusions thermogenic gas, mostly methane, exists in the lopra1/1a well in the hyaloclastites and in the more porous tuffaceous zones of the overlying lower basalt formation. chemical and isotopic data suggest that gases in the hyaloclastites were responsible for the more or less continuous seepage of gas into the original lopra-1 well since 1983. the location and the type of source rock that generated the gases observed in the lopra-1 well are still uncertain. however, recent information on the pre-volcanic, organic-rich sediments of central east greenland shows that these still have a potential for generating hydrocartw ow ay t im e (s ec ) 0 2 4 6 senw 0 50 100 150 200 km lopra-1/1a 20 km to w f a r o e s s h e l f faroe – shet land bas in westray ridge rona ridge schiehallion 13 km offseteast faroe high triassic paleocene upp er cret ace ous eocene–oligocene neogene l. cretaceous u. jurassic volcanic ? basement fig. 8. geological profile modified from spencer et al. (1999). location shown on fig. 7. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1989 90 bons (ro = 1.2%) after having been buried below 6 km of volcanic rocks. given the close connection demonstrated between central east greenland and the faroes block prior to continental break-up, it is probable that the source rock for the gases in lopra-1/1a exists below the volcanics of the faroe islands area. the possibility of migration of gas from the faroe– shetland basin towards the faroes cannot be excluded, although no evidence in support of this hypothesis has yet been found. acknowledgements the two reviewers, melody a. rooney and ger van graas, are thankfully acknowledged for their careful review of the manuscript and valuable comments which helped improve the clarity of the paper. references balling, n., kristiansen, j.i. & saxov, s. 1984: geothermal measurements from the vestmanna-1 and lopra-1 boreholes. in: berthelsen, o., noe-nygaard, a. & rasmussen, j. (eds): the deep drilling project 1980–1981 in the faeroe islands. annales societatis scientiarum faeroensis, supplementum ix, 137–147. tórshavn: føroya fródskaparfelag. berner, u. & faber, e. 1996: empirical carbon isotope/maturity relationships for gases from algal kerogens and terrigenous organic matter, based on dry, open-system pyrolysis. organic geochemistry 24, 947–955. bojesen-koefoed, j. & nytoft, h.p. 2006: petroleum geochemistry of the deepened lopra-1/1a re-entry well, faroe islands. geological survey of denmark and greenland bulletin 9, 67–77 (this volume). coleman, d.d., risatti, j.b. & schoell, m. 1981: fractionation of carbon and hydrogen isotopes by methane-oxidizing bacteria. geochimica et cosmochimica acta 45, 1033–1037. faber, e. 1987: zur isotopengeochemie gasförmiger kohlenwasserstoffe. erdöl, erdgas, kohle 103, 210–218. faber, e., gerling, p. & dumke, i. 1988: gaseous hydrocarbons of unknown origin found while drilling. organic geochemistry 13, 875–879. faber, e., whiticar, j. & gerling, p. 1999: comparison of hydrocarbons from unconventional sources: ktb, epr and bit metamorphism. geologisches jahrbuch d107, 175–194. glassley, w.e. 2006: mineralogical and thermodynamic constraints on palaeogene palaeotemperature conditions during low-grade metamorphism of basaltic lavas recovered from the lopra-1/1a deep hole, faroe islands. geological survey of denmark and greenland bulletin 9, 109–118 (this volume). jacobsen, o.s. & laier, t. 1984: analysis of gas and water samples from the vestmanna-1 and lopra-1 wells, faeroe islands. in: berthelsen, o., noe-nygaard, a. & rasmussen, j. (eds): the deep drilling project 1980–1981 in the faeroe islands. annales societatis scientiarum faeroensis, supplementum ix, 149– 155. tórshavn: føroya fródskaparfelag. konnerup-madsen, j. 2006: a reconnaissance study of fluid inclusions in fracture-filling quartz and calcite from the lopra-1/1a well, faroe islands. geological survey of denmark and greenland bulletin 9, 119–122 (this volume). laier, t. 1999: the siljan deep well – hydrocarbon gas results. geologisches jahrbuch d107, 153–163. laier, t., nytoft h.p., jørgensen, o. & isaksen, g.h. 1997: hydrocarbon traces in the tertiary of the faeroe islands. marine and petroleum geology 14, 257–266. larsen, l.m., waagstein, r., pedersen, a.k. & storey, m. 1999: trans-atlantic correlation of the palaeogene volcanic successions in the faeroe islands and east greenland. journal of the geological society (london) 156, 1081–1095. larsen, m., hamberg, l., olaussen, s., nørgård-pedersen, n. & stemmerik, l. 1999: basin evolution in southern east greenland: an outcrop analog for cretaceous–paleogene basins on the north atlantic volcanic margin. american association of petroleum geologists bulletin 88, 1236–1261. rooney, m.a., claypool, g.e. & chung, h.m. 1995: modelling gas generation using carbon isotope ratios of natural gas hydrocarbons. chemical geology 126, 219–232. scotchman, i.c., griffith, c.e., holmes, a.j. & jones, d.m. 1998: the jurassic petroleum system north and west of britain: a geochemical oil source correlation study. organic geochemistry 29, 671–700. spencer, a.m., birkeland, ø., knag, g.ø. & fredsted, r. 1999: petroleum systems of the atlantic margin of northwest europe. in: fleet, a.j. & boldy, s.a.r. (eds): petroleum geology of northwest europe: proceedings of the 5th conference, 231–246. london: geological society. waagstein, r. 2006: composite log from the lopra-1/1a well, faroe islands. geological survey of denmark and greenland bulletin 9, in pocket inside back cover (this volume). waagstein, r., hald, n., jørgensen, o., nielsen, p.h., noe-nygaard, a., rasmussen, j. & schönharting, g. 1984: deep drilling on the faeroe islands. bulletin of the geological society of denmark 32, 133–138. manuscipt received 17 june 2000; revision accepted 30 january 2001. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1990 1 geological survey of denmark and greenland bulletin 41 • 2018 review of survey activities 2017 edited by adam a. garde, ole bennike and w. stuart watt geological survey of denmark and greenland danish ministry of energy, utilities and climate 22 geological survey of denmark and greenland bulletin 41 keywords geological survey of denmark and greenland, survey organisations, current research, denmark, greenland. cover photographs from left to right 1. hyperspectral analysis of inaccessible rock faces using landor sea-based platforms has potential to become a powerful geological mapping tool, as individual mineral species can be identified with this method. example from central west greenland. 2. 3d imaging by hand-held digital photogrammetry is used in both current geological mapping and assessment of landslide risks in greenland. photograph: jonas petersen. 3. potential storage sites for unwanted brine around the lille torup gas storage facility in jylland, and seismic lines used in the location of suitable sites. 4. studying intense glaciotectonic fracturing of cretaceous chalk at stevns klint. frontispiece: facing page the geus ice-sheet field camp photographed on 20 july, 2017 above the former camp century buried in the greenland ice sheet near thule air base. chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, department of geoscience, aarhus university; minik rosing, geological museum, university of copenhagen; finn surlyk, department of geosciences and natural resource management, university of copenhagen scientific editors: adam a. garde, ole bennike and w. stuart watt editorial secretary: jane holst referees (numbers refer to first page of reviewed article): henrik friis & ida l. fabricius, dk (9); reinhard kirsch, de & michael engkilde, dk (13), theis i. sølling & jan audun, dk (17); jakob qvortrup christensen & lars kristiansen, dk (21); jakob qvortrup christensen & jette sørensen, dk (25); jørgen c. toft & arne thorshøj, dk (29); nicolaj krog larsen & anders bjørk, dk (33); andy whitham, uk & john korstgård, dk (39); steve piercey, ca & hartwig frimmel, de (43); christian rogass, de & asger k. pedersen, dk (47, 51); chris harrison & kate dickie, ca (57); asger k. pedersen, dk & ken mccaffrey, uk (63); louwrens hacquebord, nl & ole humlum, no (67); niels tvis knudsen, dk & kirsty langley, gl (71); james h. lever, us & jacob clement yde, no (75); xavier fettweis, be & horst machguth, ch (79); torben schmidt, dk & martin miles, no (83); gang lui & michael engkilde, dk (87); andreas møller & peter grønkjær, dk (91); robert tomas, ch (95); birgir v. óskarsson, is & max strunk, se (99). illustrations: stefan sølberg, adaam a. garde, ole bennike and susanne rømer layout and graphic production: jane holst and jacob lind bendtsen printer: rosendahls-schultz grafisk a/s, albertslund, denmark manuscripts received: 1 february – 28 may 2018 final versions approved: 20 february – 28 june 2018 printed: 15 august 2018 erratum: the repetition of fig. 3 in fig. 4’s position on page 73 has been corrected to show fig. 4 in the online version of this bulletin. issn (print) 1604-8156, isbn (print) 978-87-7871-500-5 issn (online) 1904-4666, isbn (online) 978-87-7871-501-2 citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 41, 104 pp. available from geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark to buy bulletin in printed form please contact bogsalg@geus.dk and at www.geus.dk/bulletin41 (open access) © de nationale geologiske undersøgelser for danmark og grønland (geus), 2018 for the full text of the geus copyright clause, please refer to www.geus.dk/bulletin 3 44 39 update of the seamless 1:500 000 scale geological map of greenland based on recent field work in the wandel sea basin, eastern north greenland. k. svennevig 43 base-metal and ree anomalies in lower palaeozoic sedimentary rocks of amundsen land, central north greenland: implications for zn-pb potential d. rosa, j.f. slack and h. falck 47 mineral mapping by hyperspectral remote sensing in west greenland using airborne, ship-based and terrestrial platforms s. salehi and s.m. thaarup 51 hyperspectral analysis of lithologies in the arctic areas with abundant lichen cover s. salehi malaysia bangladesh vietnam greenlandcanada kenya ethiopia lebanon cameroon ghana senegal zambia malawi denmark norway china 7 review of survey activities 2017 f.g. christiansen 9 diagenetic impact on reservoir sandstones of the heno formation in the ravn-3 well, danish central graben s. pedersen, r. weibel, p.n. johannessen and n.h. schovsbo 13 potential for brine storage near the gas storage facility at lille torup, northern jylland, denmark .... m.l. hjuler, m.s. andersen, c.m. nielsen, a. mathiesen, l. kristensen, n. skaarup and l.h. nielsen 17 shale fabric and organic nanoporocity in lower palaeozoic shales, bornholm, denmark l.m. henningsen, c.h. jensen, n.h. schovsbo, a.t. nielsen and g.k. pedersen 21 chalk-glacitectonite, an important lithology in former glaciated terrains covering chalk and limestone bedrock s.a.s. pedersen, p. gravesen and k. hinsby 25 sedimentological and glaciotectonic interpretation of georadar data from the margin of the vig ice-push ridge, nw sjælland, denmark c.s. andersen & p.r. jakobsen 29 miocene oil-bearing diatom ooze from the north sea e. sheldon, e.s. rasmussen, k. dybkjær, t.e. eidvin, f. riis and r. weibel 33 initial observations of the shallow geology in tannis bugt, skagerrak, denmark m.j. owen, n.h. witt, z. al-hamdani, n. nørgaard-pedersen, k.j. andresen and j.o. leth 5 dark grey indicates non-european countries where geus has ongoing or recently completed projects. orange indicates countries with geus projects described in this volume. 57 new subsurface mapping offshore southern west greenland using geophysical and geological data u. gregersen, m.s. andersen, h. nøhr-hansen, e. sheldon, t.f. kokfelt, m. olivarius, c. knudsen, k.g. jakobsen and j.s. adolfssen 63 remote geological mapping using 3d photogrammetry: an example from karrat, west greenland e.v. sørensen and p. guarnieri 67 european trading, whaling and climate history of west greenland documented by historical records, drones and marine sediments n.e. mikkelsen, a. kuijpers, s. ribeiro, m. myrup, i. seiding and a.e. lennert 71 the greenland ice sheet – snowline elevations at the end of the melt seasons from 2000 to 2017 r.s. fausto and the promice team 75 initial field activities of the camp century climate monitoring programme in greenland w. colgan, a. pedersen, d. binder, h. machguth, j. abermann and m. jayred 79 circum-greenland, ice-thickness measurements collected during promice airborne surveys in 2007, 2011 and 2015 l.s. sørensen, s.b. simonsen, r. forsberg, l. stenseng, h. skourup, s.s. kristensen and w. colgan 83 observationally constrained reconstruction of 19th to mid-20th century sea-ice extent off eastern greenland d.a.m. hallé, n.b. karlsson, a.m. solgaard and c.s. andresen 87 examining the rare-earth elements (ree) supply– demand balance for future global wind power scenarios p. kalvig and e. machacek 91 analysis of cod otolith microchemistry by continuous line transects using la-icp-ms s.h. serre, k.e. nielsen, p. fink-jensen, t.b. thomsen and k. hüssy: tulstrup and m. pedersen 95 towards a common geological data infrastructure for europe j. tulstrup and m. pedersen 99 analytical procedures for 3d mapping at the photogeological laboratory of the geological survey of denmark and greenland e.v. sørensen and m. dueholm malaysia bangladesh vietnam greenlandcanada kenya ethiopia lebanon cameroon ghana senegal zambia malawi denmark norway china 66 7 review of survey activities 2017 flemming g. christiansen deputy director 2017 was a year where we all experienced how fatal geological forces can be for society. on june 17, a tragic natural disaster hit greenland. following a huge landslide into karrat fjord, major waves flooded the settlements of nuugaatsiaq and illorsuit. four persons were lost and many houses in nuugaatsiaq were destroyed. due to the continued high risk for more landslides, the inhabitants of the two settlements have not returned and the greenlandic authorities advise against visiting the risk area. such a natural disaster together with many extremes in weather and climate conditions around the world, also in denmark and greenland, directs our focus on the consequences of the changing climate. this requires regularly updated climate models and response with suggestions for adaptation to climate change and will have strong implications for geus’ continued research and monitoring in many years to come, including further studies and monitoring of geohazards. geus will have an important role in guiding society for better and safer living conditions in both denmark and greenland. it is necessary to work with such topics and provide society and authorities with transparent information. geus has presented a new website providing key information on many different research areas and with specific information and data from monitoring programmes like www. promice.dk, celebrating its 10 years anniversary in 2017 and www.campcenturyclimate.dk that was established in 2017. it is essential that such key information builds on results of a high scientific standard and that data are well documented in international publications, and in papers and maps in our own series. this issue of geus’ review of survey activities includes 22 papers covering many different activities in denmark, greenland and internationally. seven papers are on denmark, 11 on greenland and four on other themes. activities in denmark geus works with many different – and often closely related – topics in denmark such as the use of water, energy and mineral resources, protection of nature when exploiting resources, and the impact of climate change. the economic and political implications are very significant, and making up-to-date geological and geophysical data and information available for society, authorities and industry is of great value in this context. one paper is on the diagenetic impact of upper jurassic sandstones in the ravn-3 well in the danish central graben. understanding these processes is important for the production of oil from deeply buried oil fields. another paper is on the thickness, depth and properties of reservoir rocks within the sedimentary succession close to the lille torup gas storage facility in northern jylland. to increase the volume of caverns, saline brine from the salt structure must be stored elsewhere, and one of the options could be to re-inject the brine into the subsurface. a third paper deals with the controlling factors of porosity development in palaeozoic shales and the implications for shale-gas plays in denmark. an example from bornholm, demonstrates that this is controlled by both shale fabric and organic nanoporosity, in contrast to conventional sandstone reservoirs. understanding rocks formed by glacial processes is important in denmark, e.g. for flow modelling of ground water and for handling geotechnical problems when establishing new infrastructure. one paper describes the formation of chalk-glaciotectonite with examples from former glaciated terrains covering chalk and limestone bedrock from localities on møn, sjælland and jylland. a second paper presents sedimentological and glaciotectonic interpretation of georadar data from nw sjælland, where the interior structure of ridges has been characterised. for cost-reduction reasons, there is a growing interest for oil exploration in the youngest and shallowest part of the sedimentary succession in the north sea. a paper gives a detailed description of nanofossils, microfossils and palynomorphs in a miocene oil-bearing diatom ooze from the valhall field area in the norwegian sector. a final paper is on the shallow geology of tannis bugt in skagerak where habitat mapping has been carried out in a natura2000 area. the geological history here is very complex with deformed units and possibility of pre-quaternary bedrock being exposed on the sea bottom. © 2018 geus. geological survey of denmark and greenland bulletin 41, 7–8. open access: www.geus.dk/bulletin http://www.promice.dk http://www.promice.dk http://www.campcenturyclimate.dk http://www.geus.dk/bulletin 88 activities in greenland as in previous years, there was a high level of geological and glaciological activities in greenland in 2017: both traditional studies with focus on geological mapping, the mineral and petroleum potential and monitoring and research related to climate changes and their effects. one paper discusses the implications of a new 1:100 000 map from kilen in north greenland for the seamless 1:500 000 regional map of greenland. new procedures and documentation are necessary, as many geological maps from greenland are likely to be updated in digital versions in the coming years. a second paper on north greenland describes base-metal and rare-earth elements (ree) anomalies in lower palaeozoic sedimentary rocks and discusses the implications for the zinc and lead potential. hyperspectral remote sensing has a great potential for geological mapping and exploration in greenland. a paper from west greenland provides details on mineralogical mapping of basement rocks in the nagssugtoqidian orogen using airborne, ship-based and terrestrial platforms. another paper focuses on how the signal from lichens that often cover rocks in the arctic can be distinguished from the mineralogical signal, with examples of ultramafic rocks such as kimberlite. one paper presents results from subsurface mapping offshore southern west greenland using seismic interpretation also including various geological and geochemical analysis. this paper includes new and critical information from oil exploration wells drilled in 2010 and 2011 that penetrated the deeper part of the sedimentary succession and reached underlying basement and volcanic rocks. another paper uses 3d photogrammetry for lithological mapping and structural analysis, in this case from karrat fjord in west greenland. such understanding is also important for further evaluation of risk of landslides. climate-related research and monitoring at geus provide important data for global climate models and are often based on ground-truth data from the ice and fjords or offshore. one paper compares historical records of european trading and whaling in the disko bugt region with climate data from marine sediments and show interesting examples of the physical remains from the whaling period. another paper introduces an important climate indicator – the snowline elevation that is the maximum elevation during the melt season where snow remains from the previous accumulation season. based on satellite data and direct comparison with data from promice stations, a series of snowline maps covering all of greenland from 2011 to 2017 have been constructed. with new and updated climate models, there are concerns that remains from the former us camp century base buried under the ice in the thule area might get closer to the surface within the next century and that meltwater may interact with its waste long before. geus has started a new climate-monitoring programme and has mapped the extent and depth of the debris from the base with an ice-penetrating radar survey; results from the first field season are summarised in one contribution. a second paper describes results from airborne surveys in 2007, 2011 and 2015 where the elevation of the ice surface, top of bedrock, and the variation in thickness of the greenland ice sheet through time have been measured. another paper is on the sea-ice distribution along eastern greenland and iceland. original maps by lauge koch covering the 19th and 20th century have been digitised and analysed statistically. this can be very useful for future models of sea-ice variation in a changing climate. other themes as the national geological survey, geus has a strong obligation to participate in international assessments of resources, use state-of-the-art laboratory equipment, constantly develop new analytical techniques and make the ever-increasing volume of data available with systematic quality control and updated, user-friendly databases and distribution systems. one paper presents an examination of the supply-demand balance of ree, which is important for future global wind power scenarios; especially lack of neodymium and other elements for permanent magnets may be critical. another paper describes how analysis of the microchemistry of cod otoliths (ear stones) with the la-icp-ms technique can be applied to understand the age, growth history and migration of fish stock. a third paper gives an overview of the development and organisation of a common geological data infrastructure (egdi) for europe. egdi is important for the research collaboration and geological data sharing between the member states of the eurogeosurveys. finally, the last paper presents the procedures for 3d mapping at the photogeological laboratory at geus. this is a very strong tool for geological mapping, structural analysis and evaluation of the risk of natural geohazards such as landslides. a reconnaissance study of fluid inclusions in fracture-filling quartz and calcite from the lopra-1/1a well, faroe islands 119© geus, 2006. geological survey of denmark and greenland bulletin 9, 119–122. available at: www.geus.dk/publications/bull a reconnaissance study of fluid inclusions in fracture-filling quartz and calcite from the lopra-1/1a well, faroe islands jens konnerup-madsen fracture-filling calcite and quartz from the lopra-1/1a well (at 2380 m and 3543 m depth) contains both aqueous low-salinity fluid inclusions and hydrocarbon-dominated fluid inclusions. microthermometry indicates that the aqueous fluids contain 0.2 to 1.4 equivalent weight% nacl and occasionally contain traces of hydrocarbons. homogenisation to liquid occurred between 90°c and 150°c. modelling based on these fluid inclusion observations indicates that during burial the basaltic section was subjected to temperatures of 160°c and 170°c, occasional pressures of 600–700 bars and the simultaneous percolation of aqueous and hydrocarbon fluids. these fluid conditions may also be relevant to the formation of zeolite observed in the lopra-1/1a well. keywords: basalts, faroe islands, fluid inclusions, hydrocarbons, veins, zeolites __________________________________________________________________________________________________________________________________________ geological institute, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k., denmark. e-mail: jenskm@geol.ku.dk fluid inclusions in cements or minerals filling vugs and fractures in buried sedimentary and volcanic rocks may provide important information on the chemical and physical nature and origin of mineral-precipitating fluids, on the potential interplay between migrating hydrocarbon and aqueous fluids, and on the temperatures and pressures of precipitation (e.g. bodnar 1990; jensenius & burruss 1990). a reconnaissance study was undertaken of fluid inclusions in vugand fracture-filling quartz and calcite from samples taken from the basalts penetrated by lopra1/1a. the two samples studied are from core 1 (2380 m) and sidewall core 1 (3543 m). the fluid inclusions were examined by ordinary microscopy, fluorescence microscopy and with a chaixmeca heating and freezing stage. types and setting of fluid inclusions in selected samples the samples were selected by examining about 40 thin sections taken between 2204 m and 3543 m depth in the lopra-1/1a well. only two samples, from 2380 m and 3543 m depth, were found to contain fracture-filling quartz and calcite with fluid inclusions suitable for further study. sample 2380 m (lopra-1, core 1) is a sparsely plagioclase-glomerophyric olivine-clinopyroxene basalt with almost complete alteration of plagioclase and olivine. the quartz and calcite studied occur in mm-wide veins. the veins are rimmed by chlorite, calcite and quartz that appear to have been precipitated contemporaneously. according to jørgensen (2006, this volume) the zeolites characterising this level in the core are laumontite, prehnite and pumpellyite. sample 3543 m (lopra-1a, sidewall core 1) is a nearaphyric lapilli-tuff with extensively altered plagioclase, olivine and clinopyroxene phenocrysts in a cryptocrystalline groundmass. the irregular veins contain laumontite, prehnite, calcite and rare quartz. the veins are rimmed by chlorite. again, calcite and quartz appear to have been precipitated contemporaneously, although quartz precipitation might have been slightly later. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19119 120 0 1 2 3 50 100 150 200 homogenisation temperature, °c sa lin ity , w t% n ac l group a group b data on fluid inclusions in quartz from lopra-1/1a altered basalt field for aqueous fluid inclusions in calcite types of fluid inclusions in quartz and calcite two types of fluid inclusions were observed using fluorescence and ordinary light microscopy: (1) aqueous twophase (liquid-vapour) inclusions with about 5 vol.% vapour at room temperature, and (2) one or two-phase hydrocarbon inclusions with fluorescence emission colours that vary from orange-yellow to green. with ordinary light microscopy it is difficult to distinguish between the twophase liquid-vapour hydrocarbon and aqueous inclusions, although the latter seem to be characterised by a (perhaps) slightly lower vol.% vapour than the former. no clear relative chronology between the two fluid inclusion types could be established. examples of typical morphologies and phase proportions of fluid inclusions observed in quartz are shown in fig. 1. in general, the fluid inclusions are characterised by immature morphologies and occur in irregular groupings or in curved internal planar arrangements, suggesting periods for their entrapment which do not markedly postdate the growth of the host mineral. the liquid-vapour ratio in individual groupings varied slightly, most probably and mainly as a result of necking down of the inclusions after entrapment, because liquid-only inclusions could occasionally be observed together with the aqueous two-phase liquid-vapour inclusions. all inclusions indicative of having been influenced by necking down were avoided during the heating and freezing stage work. microthermometry results on aqueous fluid inclusions the results of microthermometry of fluid inclusions in quartz and calcite are summarised in fig. 2. fluid inclusions in quartz incipient melting of ice was observed at temperatures around –32°c, indicating the presence of additional ions such as ca2+, mg2+ and/or fe2+ in solution rather than chlorides of na+ and/or k+ (konnerup-madsen 1979). final melting temperatures were observed in the range –0.1°c to –0.9°c, corresponding to salinities from 0.167 to 1.49 equivalent weight% nacl, respectively (average: 0.62 equivalent weight% nacl) (bodnar et al. 1989), but with no clear difference between the two samples. temperatures of homogenisation occurred between 94°c and 150°c and bimodality in temperature is suggested from the data (see fig. 2, groups a and b). group a and group b inclusions gave average homogenisation temperatures of 108°c and 141°c, respectively. group b inclusions in quartz showed in three cases clear indications (ragged outline of meniscus between vapour and liquid) of the formation of a clathrate hydrate after initial ice melting, indicating the presence of trace amounts of volatiles such as hydrocarbons in the entrapped group b fluids. however, although no temperature of dissolution of the hydrate could be obtained and hence the identity of the volatile component could not be established, its formation suggests that the higher temperatures of homogenisation obtained for group b inclusions may reflect trace concentrations of hydrocarbons in the vapour phase of these inclusions. fig. 1. examples of typical morphologies of aqueous liquid-vapour fluid inclusions in quartz from core 1 (2380 m) from lopra-1/1a. fig. 2. salinity versus liquid homogenisation temperatures of aqueous inclusions in quartz from lopra-1/1a. 10 µm geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19120 121 fluid inclusions in calcite only very few measurements were made on inclusions in calcite, as most inclusions occurred along well-defined healed fracture-planes so are secondary in origin. final melting temperatures varied between –0.2°c and –0.7°c, corresponding to salinities of 0.33 to 1.16 equivalent weight% nacl. homogenisation temperatures varied from 101°c to 186°c (fig. 2). however, the higher homogenisation temperatures might conceivably reflect partial decrepitation (stretching) of the inclusions during heating (e.g. bodnar & bethke 1984). two-phase (liquid-vapour) hydrocarbon inclusions were observed in fluorescence microscopy in both calcite and quartz. the abundance of hydrocarbon inclusions appears to be relatively higher in calcite and only very few were observed in quartz. the emission colours, from orangeyellow to green, may be interpreted roughly in terms of compositions corresponding to api gravities of 20–35 (lang & gelfand 1985). no successful heating and cooling runs were, however, obtained on the hydrocarbon inclusions in the two samples. interpretation of fluid inclusion data an interpretation in terms of pressures and temperatures for entrapment of the fluid inclusions in quartz is shown in fig. 3. isochores corresponding to group a and b inclusions (brown 1989) in quartz are shown in fig. 3, assuming them to be pure aqueous fluids with salinities as indicated by the final ice melting temperatures (fig. 2). no indications of the entrapment of boiling aqueous fluids were observed during this study and the homogenisation temperatures for the fluid inclusions observed are thus considered to be minimum temperatures of fluid entrapment and host mineral formation. a comparison with geothermal gradients of 20°c/100 bars and 20°c/226 bars that are considered relevant for lopra-1/1a and that reflect hydrostatic and lithostatic conditions, respectively, has been made in fig. 3. if hydrostatic conditions prevailed, group a inclusions would indicate entrapment at around 140°c at pressures of around 600 bars. microthermometry indicated that group b inclusions may contain traces of hydrocarbons and the isochores shown in fig. 3 are therefore not strictly applicable because they assume an aqueous-only composition. as trace concentrations of hydrocarbons are present in group b inclusions, pressures at homogenisation will be considerably higher than indicated by the isochores drawn in fig. 3. the presence of only a few parts per thousand methane in solution would shift homogenisation pressures to values of 400–600 bars at the observed temperatures of homogenisation (hanor 1980). the actual isochoric path for group b inclusions should therefore be shifted to a setting essentially parallel to that shown but starting at the bubble-point curve for the actual aqueous-hydrocarbon system at around 400 bars (fig. 3, point a). if this interpretation is valid, both groups of inclusions in quartz indicate minimum entrapment of fluids slightly different in composition at conditions of about 400 bars and 140°c. assuming hydrostatic conditions, probable entrapment of both group a and b aqueous fluids low in salts (average 0.61 equivalent weight% nacl) and containing occasional traces of hydrocarbons occurred at around 600–700 bars at temperatures of 160°c to 170°c. however, more data would be needed to substantiate this conclusion. concluding remarks although it is of a reconnaissance nature, the present study of fluid inclusions in fracture-filling quartz and calcite indicates that the basaltic sections represented by the samples examined were subjected to temperatures of 160°c to 170°c and pressures of 600–700 bars at stages during their burial. during these burial conditions, precipitation of quartz and calcite in fractures (and vugs?) occurred in the presence of low-salinity aqueous fluids containing occasional traces of hydrocarbons. similar p–t-fluid-char0 500 1000 1500 2000 0 100 200 300 temperature, °c pr es su re , ba rs hydrostatic gradient 20°c/226 bars hydrostatic gradient 20°c/100 bars bubble-point curve for h 2 o–0.2 mole % ch 4 a group a group b isochore for aqueous inclusions fig. 3. pressure-temperature diagram with isochores for groups a and b inclusions in quartz from lopra-1/1a. the open and filled circles show pressure and temperature at homogenisation for pure aqueous and aqueous-0.2 mole%ch4 fluids in group b inclusions, respectively. bubble-point curve from hanor (1980). see text for further comments. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19121 122 acteristics may also be of relevance to the formation of e.g. zeolites in these rocks. although no clear evidence for the simultaneous existence and migration of hydrocarbon and aqueous fluids was observed, such simultaneity is suggested by the occasional presence of hydrocarbons in the entrapped aqueous fluids and the hydrocarbon-dominated inclusions observed especially in calcite. references bodnar, r.j. 1990: petroleum migration in the miocene monterey formation, california, usa: constraints from fluid-inclusion studies. mineralogical magazine 54, 295–304. bodnar, r.j. & bethke, p.m. 1984: systematic stretching of fluid inclusions. fluorite and sphalerite at one atmosphere confining pressure. economic geology 79, 141–146. bodnar, r.j., sterner, s.m. & hall, d.l. 1989: salty: a fortran program to calculate compositions of fluid inclusions in the system nacl-kcl-h2o. computers & geosciences 15, 19–41. brown, p.e. 1989: flincor: a microcomputer program for the reduction and investigation of fluid inclusion data. american mineralogist 74, 1390–1393. hanor, j.s. 1980: dissolved methane in sedimentary brines: potential effect on the pvt properties of fluid inclusions. economic geology 75, 603–617. jensenius, j. & burruss, r.c. 1990: hydrocarbon-water interactions during brine migration: evidence from the composition of hydrocarbon inclusions in calcite from danish north sea oil fields. geochemica cosmochemica acta 54, 705–713. jørgensen, o. 2006: the regional distribution of zeolites in the basalts of the faroe islands and the significance of zeolites as palaeotemperature indicators. geological survey of denmark and greenland bulletin 9, 123–156 (this volume). konnerup-madsen, j. 1979: fluid inclusions in quartz from deepseated granitic intrusions, south norway. lithos 12, 13–23. lang, w.h. & gelfand, j.c. 1985: the evaluation of shallow potential in a deep field wildcat. log analyst 26, 13–22. manuscipt received 15 december 1999; revision accepted 29 june 2001. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19122 geological survey of denmark and greenland bulletin 26, 2012, 69-72 69 methane and possible gas hydrates in the disko bugt region, central west greenland naja mikkelsen, troels laier, tove nielsen, antoon kuijpers and niels nørgaard-pedersen current climate models predict an annual temperature increase in the arctic between 4° and 6°c by the end of the 21st century with widespread impact on the arctic environment. warming will lead to thawing of the widespread, permanently frozen, high-latitude peat-lands and to degradation of marine gas hydrates, both of which may increase the rate of methane release to the atmosphere. this will influence global climate as methane is a potent greenhouse gas with a large global warming potential. marine gas hydrates are found worldwide on continental margins and frequently occur in the arctic. interpretation of seismic profiles has also indicated their presence in the disko bugt region in western greenland. in june 2011 a scientific cruise was undertaken in the disko bugt region (fig. 1) to investigate the occurrence of methane and possible gas hydrates in the region. the cruise was part of a multidisciplinary scientific project impact on permafrost, gas hydrates and periglacial processes following climate changes in greenland (permagas). the project studies the impact of global climate warming on permafrost and gas hydrates in the disko bugt region. the aim of the project is to link marine and terrestrial occurrences of gas emissions. marine gas hydrates gas hydrate is a crystalline solid consisting of gas molecules, usually methane, with each gas molecule surrounded by a frame of water molecules. marine gas hydrates form under high pressure and low temperature in sediments below the seabed (fig. 2). depending on the bottom water temperature methane hydrate is typically stable in sea-floor sediments on the continental slope, but in high-latitude regions with low bottom water temperatures, the top of the gas hydrate stability zone may occur at shallower depths. gas hydrates are a potential energy resource as well as a potential risk for geohazards and the safe exploitation of sea bed resources (kvenvolden & rogers 2005). the worldwide amount of carbon bound in gas hydrates is conservatively estimated to total twice the amount of carbon found in all fig. 1. map of the disko bugt region. the black dots show core sites in egedesminde dyb (a, b), on the shelf off aasiaat (c), in southern vaigat (d) and off jakobshavn isfjord (e). methane was encountered at sediment sites a and b, and traces of methane occurred at site c. at sites d and e, where pockmarks have been mapped during previous surveys, evidence of upwelling freshwater was found. contours: 100, 200 and 500 m. the red line that crosses the core location at site c shows the position of the seismic profile in fig. 3. fig. 2. phase diagram showing the boundary between free methane (no colour) and gas hydrate (blue). vaigat 69° 54°w 25 km 56° 68°30´n 200 500 50 0 disko disko bugt jakobshavn isfjord aasiaat a b c e d greenland 20 0 20 0 2 4 6 8 10 temperature (°c) 5 10 150 p re ss u re ( m p a) site b egedesminde dyb gas hydrate the general water depth in disko bugt is 200–400 m, corresponding to c. 2–4 mpa methane + water © 2012 geus. geological survey of denmark and greenland bulletin 26, 69–72. open access: www.geus.dk/publications/bull 7070 known fossil fuels on earth, and methane bound in hydrates amounts to approximately 3000 times the volume of methane in the atmosphere. in a warming world, methane from the dissociation of large and dynamic gas hydrate reservoirs therefore has the potential to influence oceanic and atmospheric carbon pools and thus influence global climate. gas hydrate may be recognised on seismic profiles by the presence of a so-called bottom-simulating reflector that marks the base of the gas hydrate stability zone (mackay et al. 1994). the reflector is caused by the impedance contrast between the solid gas hydrate layer and free gas accumulations beneath. however, bottom-simulating reflectors are also found in areas without gas hydrates. the disko bugt region since the discovery of extensive oil seeps north of disko in 1992 (christiansen et al. 1996), marine geologists’ interest in the disko bugt region has increased significantly. however, little is known about the possible existence of gas hydrates on the continental margin, offshore west greenland. the presence of pockmarks in disko bugt (weinrebe et al. 2008) provides evidence of upwelling gas or fluid from the sea bottom. the bottom water temperature is c. 3°c in the disko bugt region and gas hydrates can be expected to occur at water depths exceeding 400 m, provided that methane occurs in high concentrations. during collection of a piston core in central disko bugt high gas content was demonstrated by the sudden escape of large amounts of strongly expanding gas that disrupted the sediment (kuijpers et al. 2001). in addition, bottom-simulating reflectors have been observed on a number of seismic profiles from the area (fig. 3). material and methods during the cruise, up to 6 m long gravity cores and up to 2 m long cores, taken with a max planck institute rumohr lot corer, were retrieved from five sites above and within the gas hydrate stability zone. the cores were sub-sampled for analysis of sediments and pore fluids (fig. 4; nielsen et al. 2011). the sediment cores were subjected to a number of geochemifig. 3. reflection seismic profile from site c (fig. 1). the bottom-simulating reflector (yellow) at 75 msec two-way travel time below the seabed may indicate the occurrence of an up to 70 m thick gas hydrate zone. pockmarks and seabed mounds (red) overlying faults in the shallow sub-seabed unit may be caused by seepage of free gas from beneath the gas hydrate zone. the seismic profile is part of released data acquired for the company nunaoil in 1998. fig. 4. plastic liners with 10 cm diameter cores sampled for pore water. the samples were analysed during the cruise for concentrations of methane, sulphate and dissolved sulphide. the rhizon samplers were inserted into pre-drilled holes in the gravity core sections and pore water extracted by applying vacuum to the sampler when pulling the syringe piston. fig. 5. total inorganic carbon (tic), total sulphur (ts) and total organic carbon (toc) in a sediment core from site b in egedesminde dyb (after kuijpers et al. 2001). 1 km 0.5 1.0 tw o -w ay t ra ve l ti m e (s ec ) sw ne tic (%) ts (%) toc (%) 0 1 2 3 0 300 600 900 d ep th ( cm b el o w c o re t o p ) 71 cal analyses, including measurements of methane concentrations and concentration of pore-water solutes (particularly sulphate), which aimed at providing data that could confirm the presence of gas hydrates. geochemical results and discussion in the cores collected in egedesminde dyb at sites a and b (fig. 1), which are situated within the gas hydrate stability zone, high pore-water methane concentrations were noted. however, the maximum methane concentration measured (c. 16 mm) is much below what is expected for methane saturation at 800 m depth in the egedesminde dyb (147 mm; yamamoto et al. 1976). this is probably due to partial degassing during core retrieval. a pressure core barrel was not available during the cruise and the large drop in pressure during retrieval of the sediment cores would inevitably lead to loss of methane. therefore it was not possible to prove the existence of small amounts of methane hydrate that may have formed as a result of in situ methane supersaturation. we did not observe any diagnostic features of hydrates either, such as trends in the chloride concentration or soupy sediment textures. the methane is most likely microbial in origin and formed in situ as a result of organic matter degradation below the sulphate zone c. 0.5 m below the sea floor. this assumption is supported by the high content of organic matter in the sediment (total organic carbon = 1.9–2.3%; fig. 5). methane production is also promoted by the relatively high sedimentation rate in the area, 0.4–0.5 cm/year at site b (moros et al. fig. 6. pore-water concentration profiles of dissolved methane, sulphate and chloride from sediment cores retrieved from egedesminde dyb (sites a (842 m) and b (865 m)), off aasiaat (site c, 544 m) and from vaigat (site d, 469 m). yellow symbols: short rumohr lot cores, black symbols: gravity cores. concentrations are in mm (millimoles per litre). 440 480 520 560 gravity core rumohr lot core 440 480 520 560 cl mm 440 480 520 560 440 480 520 560 0 10 20 30 0 10 20 30 0 10 20 30 so4 mm 0 10 20 30 0 100 200 300 400 500 600 0 0.1 0.2 0 100 200 300 400 500 600 0 5 10 15 20 0 100 200 300 400 500 600 d ep th ( cm b el o w c o re t o p ) 0 5 10 15 20 0 100 200 300 400 500 600 ch4 mm 0 5 10 15 20 site a site b site c site d sulphate-methane transition zone sulphate-methane transition zone 7272 2006) that allows for a high proportion of easily degradable organic matter to enter the zone of methanogenesis (henrichs & reeburgh 1987). at site c in the area west of aasiaat the much lower methane concentration than at sites a and b may be explained by the high sulphate concentration (figs 1, 6), which generally excludes the presence of methane (iversen & jørgensen 1985). still the concentration of methane is significantly above background values for other sulphate pore-water concentrations in the disko bugt region. this may suggest upward migration of methane from gas hydrates as indicated by seismic data from the area (fig. 3). the slight decrease in pore-water chloride concentrations with depth (fig. 6) may further sustain the assumption of an upward migration of fluids depleted in chloride from below. pore-water sulphate in sediment cores from site d located in the southern end of the strait vaigat is almost exhausted at approximately 5 m below the sea floor (figs 1, 6). however, the low sulphate concentration is presumably not entirely due to in situ microbial sulphate reduction, as the decrease in pore-water chloride concentration with depth indicates a considerable contribution of freshwater from submarine groundwater discharge (fig. 6). additional field work was conducted during the cruise in an area off the mouth of jakobshavn isfjord (site e). pockmarks observed during an earlier multibeam survey in that area (weinrebe et al. 2008) were suspected to have formed due to upward gas migration (hovland & svendsen 2006). however, the absence of methane together with the sediment texture observed in sediment cores from the area indicate that the pockmarks form from upwelling water and not from gas seepage. concluding remarks the geochemical data obtained as a result of the 2011 cruise to the disko bugt region indicate that gas hydrates may occur in the region. further investigation of the possible gas hydrates will continue during a new cruise in the area in 2012. acknowledgements the 2011 cruise was supported by a grant from geocenter denmark to the project permagas. the success of the cruise depended very much on the operational skills of the captain and crew of the r/v paamiut and the smooth co-operation during the cruise with a scientific team from the greenland institute of natural resources. references christiansen, f. g., bate, k.j., dam, g., marcussen, c. & pulvertaft, t.c.r. 1996: continued geophysical and petroleum geological activities in west greenland in 1995 and the start of onshore exploration. bulletin grønlands geologiske undersøgelse 172, 15–21. henrichs, s.m. & reeburgh, w.s. 1987: anaerobic mineralization of marine sediments organic matter: rates and role of anaerobic processes in the oceanic carbon economy. geomicrobiology journal 5, 191–237. hovland, m. & svendsen, h. 2006: submarine pingoes: indicators of shallow gas hydrates in a pockmark at nyegga, norwegian sea. marine geology 228, 15–23. iversen, n. & jørgensen, b.b. 1985: anaerobic methane oxidation rates at the sulphate–methane transition in marine sediments from kattegat and skagerrak (denmark). limnology oceanography 30, 944–955. kvenvolden  k.a. &  rogers  b.w.  2005:  gaia’s breath – global methane exhalations. marine and petroleum geology 22, 579–590. kuijpers, a., lloyd, j.m., jensen, j.b., endler, r., moros, m., park, l.a., schulz, b., jensen, k.g. & laier, t. 2001: late quaternary circulation changes and sedimentation in disko bugt and adjacent fjords, central west greenland. geology of greenland survey bulletin 189, 41–47. mackay, m.e., jarrard, r.d., westbrook, g.k. & hyndman, r.d. 1994: origin of bottom-simulating reflectors: geophysical evidence from the cascadia accretionary prism. geology 22, 459–462, http://dx.doi. org/10.1130/0091-7613(1994)022%3c0459:oobsrg%3e2.3.co;2 moros, m., jensen, k.g. & kuijpers, a. 2006: midto late-holocene hydrological and climatic variability in disko bugt, central west greenland. the holocene 16, 357–367. nielsen, t., laier, t., mikkelsen, n. & kristensen, j.b. 2011: permagas project: sampling gas hydrates in the disko bay area. cruise report – r/v paamiut 20 to 26 june 2011. danmarks og grønlands geologiske undersøgelser rapport 2011/105, 40 pp. weinrebe, w., kuijpers, a., klaucke, i., fink, m., jensen, j.b. & mikkelsen, n. 2008: high-resolution bathymetry of disko bay and ilulissat icefjord, west greenland. american geophysical union fall meeting, san fransisco, california, 15–19 december 2008 (abstract c31e0563). yamamoto, s., alcauskas, j.b. & crozier, t.e. 1976: solubility of methane in distilled water and seawater. journal of chemical and engineering data 21, 78–80. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: nm@geus.dk geological survey of denmark and greenland bulletin 23, 2011, 41–44 41 since 1991, a number of european satellites have acquired data of the earth’s surface for environmental monitoring. in general, a satellite will orbit the earth in about 1½ hours and it takes 35 days before an ers or envisat satellite repeats radar scanning of the same position. for younger generations of satellites, such as radarsat and terra, the scanning repeat interval has decreased to 24 and 11 days, respectively, so that hundreds of radar scenes of the same place, produced over the past c. 20 years, are now available. persistent scatterer interferometry (psi) is a remote-sensing technique for measuring and monitoring land deformation that uses these radar scenes (ferreti et al. 2001). the technique can be used to assess natural ground movements and displacement of man-made constructions. over the next three or more years the geological survey of denmark and greenland (geus) will participate in three satellite monitoring projects conducted under the auspices of the european union. they are all funded under the global monitoring environment system (gmes). the first project, which is the subject of this paper, is named terrafirma and started in 2003 as a european space agency gmes service project. the second project, funded by the 7th framework programme, is named subcoast, and it will monitor subsidence in coastal areas. in denmark, it will concentrate on the southern part of the island of lolland. finally a large monitoring project named pangeo with similar funding has begun early in 2011. twenty-seven european geological surveys participate in pangeo that focuses on ground movements in urban areas. two cities in each of the participating countries are selected as targets for psi analysis. concerns about the effects of global climate changes are the main motivation for the gmes support to the satellite monitoring projects. the terrafirma project was extended with an additional three years of research and development and will continue until 2012. the project has five themes: (1) tectonic movements, (2) hydrological conditions, (3) flooding, (4) subsidence in abandoned mining areas and (5) wide-area satellite scanning. geus is involved in the flooding theme dealing with the increased risks of flooding of the low-lying areas in south-west jylland, adjacent to the danish wadden sea (vadehavet). the main environmental and constructional concerns are the dykes that protect the low land areas along the coast of vadehavet. in order to improve risk management detection of terrain changes in southern denmark using persistent scatterer interferometry stig a. schack pedersen, geraint cooksley, marc gaset and peter roll jakobsen 10°e 55°n germany 40 km jylland fyn rømø fig. 3 figs 4, 5 fig. 2 fig. 1. map of south-western denmark showing the location of the investigated areas. the blue frames show areas for which data from the descending track 337 and the ascending track 403 were psi processed. freshwater deposits marine sand tidal deposits aeolian sand till meltwater sand meltwater clay outwash-plain sand 20 km postglacial deposits extramarginal deposits glacial deposits 55°n 8°30´e germany rømø fanø bugt vadehavet fig. 2. map of the surface deposits in the south-western part of jylland. the region is dominated by glacial deposits of saalian age. the glacial landscape is intersected by outwash plain deposits of weichselian age. holocene tidal deposits and recent aeolian deposits are found in the westernmost part of the map. simplified from pedersen (1989). © geus, 2011. geological survey of denmark and greenland bulletin 23, 41–44. open access: www.geus.dk/publications/bull 4242 and mitigation, it is also important to identify areas of land subsidence that can be caused by geological processes and by man-made impact. the danish flood theme site the investigated area is located in the south-western part of jylland, denmark (fig. 1) that is prone to flooding, when spring tides coincide with stormy weather. after identification of the appropriate satellite tracks for the area, geus provided position data for the partner altamira information responsible for the calibration and processing of the satellite data. the coverage by the two satellites ers and envisat and the psi processing areas are shown in fig. 1. the satellite line coverage consists of both descending and ascending tracks, which in broad terms means scanning by a satellite moving both from north to south and from south to north. geus’ role in the project is to contribute with a geological and geomorphological analysis of the region (fig. 2). furthermore geus provides geographical information system (gis) data and interpretation of the data based on a geo-scientific understanding of the region. an important geus contribution to the project was the conclusion of the geological map of rømø (jakobsen 2011). rømø is located in the centre of the area covered by the satellite imagery. psi data and gis processing of satellite data a preliminary example of psi-processed data is shown in fig. 3. the orange and red pixels in the satellite image represent places where elevation changes have been detected. the psi data have been analysed using the program arcgis. the first step of this is to calibrate the data to fit relevant intervals. in the second step, the point data are statistically treated to cover the geographical area by average figures in equiva55°20´n rømø sylt riberiberibe esbjerg 8°30´e 10 km c. 6 c. 4 c. 2 0 subsidence (mm/yr) fig. 3. map of south-western jylland showing psi-processed radar data from the ers descending track 337. subsidence is seen along the causeway to sylt and along the railway line east of esbjerg. for location see fig. 1. 43 lent pixels. finally the data are compared with other terrain data, such as topographical maps, orthophotographs or terrain models based on gradient variation or geological maps. interpretation of the persistent scatterer interferometry data the obvious interpretations that can be made from the preliminary persistent scatterer interferometry processing of the satellite data are terrain movements related to man-made constructions. from this it is evident that the causeway connecting the german island of sylt to the mainland is subject to subsidence, in particular at its eastern part (fig. 3). there is also marked subsidence along the railway line east of esbjerg, and some of the bridges that are built across small streams are settling (jakobsen 2008). detailed analysis of a subsiding area: the ribe case the town of ribe and the surrounding area are described as an example of a detailed analysis using arcgis programming (figs 3–5). on the map of ribe and its surrounding, a marked subsidence is seen in the western part of the old town (yellow to red colours in fig. 4). the subsidence rate is –1 ± 0.5 mm/year, and appears to increase to c. –2–3 mm/ year on the slopes close to the stream ribe å and a smaller stream south of the town (figs 4, 5). the eastern part of ribe appears to be comparatively stable with no significant movement recorded. there is no obvious geological explanation for the subsidence in the western part of old ribe. the town is located on the eastern part of an island of marine sand surrounded by meltwater sand deposited during the saalian. during the weichselian, the ribe island became separated from the saalian deposits to the north and south by eroding rivers flowing westwards from an ice margin 30 km to the east. glaciofluvial sand and gravel were deposited by the rivers. in the holocene, ribe was situated at the boundary between an isolated sandy hill that formed an erosional remnant of the former glacial landscape surrounded by tidal flats to the west and freshwater deposits to the east. none of these geological features can explain the subsidence of western ribe. however, ribe is an old town with a long and famous historical record going back to the early part of the viking period. around ad 1100 the town was a centre for trade with a welldeveloped harbour, and was favoured with privileges given by the danish kings. during historical time, the estuary west of ribe silted up due to accumulation of tidal deposits, and ribe’s value as a merchant town decreased. several destructive events, including serious flooding, also affected ribe, therefore the town has been rebuilt several times on the rubbles of former buildings. dump and fill deposits up to 6 m thick lie beneath the present-day centre of old ribe. 2 km 8°45´e 55°20´n ribe å 0.5 to 1.5 −0.5 to 0.5 −3 to −1.5 −14.2 to −3 change (mm/yr) −1.5 to −0.5 bugt fanø 2 km ribe å ribe postglacial deposits extramarginal deposits glacial deposits freshwater sand freshwater clay freshwater gyttja freshwater peat marine sand tidal deposits aeolian sand outwash-plain sand meltwater sand till town water 55°20´n 8°45´e 5 10 5 fig. 5. detailed geological map of ribe and the ribe å area. contour interval 2.5 m. for location see fig. 1. fig. 4. image of the area around ribe (centre of image). the coloured pixels show places where elevation changes have been detected. the radar data have been psi processed in a 500 × 500 m grid with average values of the psi points representing the mean values of the vertical movements (differences in vertical displacement). the concentration of yellow pixels in the western part of ribe town is interpreted as subsidence due to urban fill. for location see fig. 1. 4444 therefore we interpret the subsidence of its western part as an effect of consolidation of the historical fill below younger buildings and constructions. however, we regard the more significant displacement on the slopes towards the river north and south of the centre as an effect of additional compaction of soft organic-rich sediments found at the transition between the freshwater drainage system and the tidal environment. detailed analysis of elevation changes: the rømø dune field one of the main geological features, expected to show up in the psi-processed data, is the concealed tønder graben (lykke-andersen 1995: gravesen et al. 2004). however, we could not identify this structure in the data. on the contrary, it appears that the data from the ascending satellite ers track 401 indicate a small regional uplift. the data central to this problem are being analysed further, but we note that some points along the west coast of rømø indicate uplift (figs 6). an uplift rate of 2 mm/year is indicated from the psi data along a row of points that coincide with the outermost dunes along the flat sandy beach plain. there are no houses or constructions in this area, so we interpret the points to represent crests of recent dunes. thus the small elevation change is caused by sand accumulation on the crest of the dunes, and the magnitude of accumulation, 2 mm/year, is a realistic figure for aeolian deposition in this area. conclusions psi-processed satellite data from vadehavet (the danish wadden sea) in south-west denmark have been analysed using arcgis and the first results indicate that no subsurface movements can be detected. a number of constructions and urban areas are subject to minor subsidence, in the order of 2–6 mm/year. a preliminary interpretation of elevation change data from the west coast of rømø implies that accumulation of dune sand is the reason for movements of around 2 mm/year. we find that psi processing of satellite data is a powerful tool for detecting elevation changes. references ferreti, a., prati, c. & rocca, f., 2001: permanent scatterers in sar interferometry. iee transactions on geoscience and remote sensing 39, 8–20. gravesen, p., jakobsen, p.r., binderup, m. & rasmussen, e.s. 2004: geologisk set: det sydlige jylland, 188 pp. copenhagen: skovog naturstyrelsen. jakobsen, p.r. 2008: geological evaluation of observed vertical terrain movements in the esbjerg test area. a contribution to the absrate/ terrafirma project. danmarks og grønlands geologiske undersøgelse rapport 2008/18, 11 pp. jakobsen, p.r. 2011: geological map of denmark, 1:50 000, 1112 iii, rømø og mandø. copenhagen: geological survey of denmark and greenland. lykke-andersen, h. 1995: neotektonik i danmark. in: nielsen, o.b. (ed.): danmarks geologi fra kridt til i dag, 19–30. århus: geologisk institut, aarhus universitet. pedersen, s.a.s. 1989: quaternary geological map of denmark, 1:200 000, map sheet 3. copenhagen: geological survey of denmark. 55°10´n 8°30´e 2 km rømø fig. 6. beach plain and sand dunes on the west coast of the island of rømø. the growing dunes are identified by the persistent scatterer interferometry technique as an area with positive elevation change. inset: close up of the island of rømø based on data from the ers satellite track 401. the scattered blue points along the west coast of the island indicate an increasing elevation of c. 2 mm/year, caused by sand eroded from the beach and deposited on the dunes. authors’ addresses s.a.s.p. & p.r.j., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: sasp@geus.dk g.c. & m.g., altamira information, còrsega 381-387, e-08037 barcelona, spain. mineralogical and thermodynamic constraints on palaeogene palaeotemperature conditions during low-grade metamorphism of basaltic lavas recovered from the lopra-1/1a deep hole, faroe islands 109 mineralogical and thermodynamic constraints on palaeogene palaeotemperature conditions during low-grade metamorphism of basaltic lavas recovered from the lopra-1/1a deep hole, faroe islands william e. glassley the sequene of secondary minerals that are reported for the lopra-1/1a well records progressive zeolite facies to prehnite–pumpellyite-facies mineral progressions consistent with those of other wellstudied hydrothermally altered rock sequences. detailed comparison of the calc–silicate (zeolites and prehnite) mineral distributions of the lopra-1/1a sequence with those from other regions indicates that this sequence exhibits consistently longer down-hole intervals for secondary mineral species than reported elsewhere. when compared to measured down-hole temperatures reported in other hydrothermally altered regions, the results suggest that the lopra-1/1a mineral progression formed under conditions typical of low temperature hydrothermal systems that form shortly after eruption of thick basaltic piles. maximum temperatures achieved at the 3500 m level of the well were at or below 200°c. the implied geothermal gradient was less than 50°c/km. an analysis of prehnite – fluid composition relationships was also conducted in order to determine if results compatible with the paragenetic sequence study could be obtained from thermodynamic constraints. in this case, the limiting temperature for prehnite formation in equilibrium with albite–quartz–calcite–laumontite (the mineral assemblage at the bottom of the hole) was determined for a range of fluid compositions. the resulting calculations suggest temperatures of formation of prehnite in the range of 140°c to 205°c, a conclusion which is broadly consistent with those reached from study of the paragenetic relationships. comparison of these results with other studies of palaeogeothermal gradients of the north atlantic margins suggests a consistent pattern in which relatively low geothermal gradients persisted in the palaeogene rift basin. keywords: north atlantic volcanic province, thermal history, geothermal gradients, low temperature metamorphism, fluid-rock interaction, reactive transport, zeolites, prehnite-pumpellyite _______________________________________________________________________________________________ lawrence livermore national laboratory, livermore, california 94550, usa. e-mail: glassley1@llnl.gov minerals that crystallise from basaltic lavas are unstable with respect to a wide range of hydrous silicates and carbonates when subjected to low temperature conditions (< 300°c) in the presence of h2oand co2-bearing fluids. recrystallisation of basaltic rocks under these physical and chemical conditions results in the development of minerals that characterise the zeolite, prehnite–pumpellyite and greenschist facies. it has been well-documented that the basalts of the east greenland – faroe islands province record extensive development of minerals characteristic of the zeolite and lower prehnite–pumpellyite facies (jørgensen 1984, 1997; neuhoff et al. 1997; larsen et al. 1999). what remains unclear is the temperature history recorded by these mineral assemblages. generally, under the lowest temperature conditions, clays, zeolites and hydrous fe–mg silicates form, giving way to less hydrated minerals at higher temperatures. often this progression is recorded by the presence of a © geus, 2006. geological survey of denmark and greenland bulletin 9, 109–118. available at: www.geus.dk/publications/bull geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19109 110 complex sequence of zeolite minerals that have increasingly smaller amounts of molecular water bound in their structures (bird et al. 1984; neuhoff & bird 2001). in principle, therefore, zeolitic and related minerals can be sensitive indicators of temperature conditions. this temperature sensitivity is complicated by the equally important sensitivity of the zeolites to the composition of coexisting fluids. the thermodynamic properties of the zeolites are affected by substitution between the alkali metals, particularly na, k and ca, and al–si exchange (e.g. neuhoff et al. 1997, 2002, 2003, 2004). the stability fields of the zeolites are also sensitive to the ratio of calcium activity to hydrogen ion activity (i.e. [ca++]/[h+]2) in the coexisting fluid phase (e.g. surdam 1973; bird et al. 1984). hence, fluid chemistry has a strong influence on both the mineral compositions that develop and the specific mineral phases that form during low temperature recrystallisation. the purpose of this paper is to define likely bounds for bottom-hole temperatures and the likely geothermal gradient active at the time of mineral development, based on paragenetic relationships and thermodynamic constraints, taking into account the effects of fluid chemistry. detailed descriptions of the locations, mineralogies and geological settings for the lopra-1/1a and vestmanna-1 boreholes are presented in other chapters in this book and are only summarised here. geology the basalts of the faroe islands were erupted subaerially onto continental crust during opening of the northern north atlantic. the basalts have been divided informally into an upper, a middle and a lower formation. the lower basaltic sequence is more than 3000 m thick (established on the basis of field exposure and the lopra-1/1a drilling programme), and ranges in age from c. 58.8 to 56.5 ma (waagstein et al. 2002). the overlying basalts and sediments (some of the sediments are coal-bearing) are more than 2000 m thick and were erupted between c. 56 and 55.5 ma (larsen et al. 1999). recrystallisation of the lavas took place during subsequent burial, leading to the development of a wide range of zeolites and associated calc–silicate minerals (jørgensen 1984, 1997). the argument that the secondary mineral development results from burial metamorphism, rather than significant tectonic stacking or folding, is based on the relatively flat-lying nature of the basaltic flows and the absence of any kinematic fabric. methods compiled published data published data from active hydrothermal systems where temperatures and mineral associations are recorded, provide the most direct evidence of the conditions under which specific mineral assemblages occur. for this reason, published data from a variety of drilled hydrothermal systems with depths less than 4000 m were analysed to identify temperature constraints that would apply to the mineral associations reported for samples from the lopra-1/1a drilling programme (jørgensen 1984, 1997). the reported lopra-1/1a assemblages were confirmed by the author during independent examination of thin sections. the best available data that correlate downhole temperatures, depth and mineral occurrences are from geothermal systems in iceland (kristmannsdóttir & tomassón 1976), japan (seki et al. 1969; boles 1981), cerro prieto (bird et al. 1984), wairakei (steiner 1977) and toa baja (cho 1991). the reports from iceland and japan discuss secondary mineral development related to alteration of basaltic rocks, which most closely correspond to the lopra1/1a sequence. the cerro prieto locality consists of sedimentary rocks (sandstones, siltstones and mudstones) that are predominately composed of quartz and feldspars. the wairakei and toa baja localities consist of volcanic and volcanoclastic rocks and their associated clastic derivatives. the wairakei rocks are primarily rhyolitic and the toa baja rocks primarily andesitic. this suite of rock types spans the entire range from basalts through andesites to rhyolites, thus encompassing silica-poor to silica-rich compositions with varying abundances of alkali metals. on a whole-rock basis, then, the compositional range from these reported systems bounds that of the faroe island basalts considered here. the different tectonic settings represented by these systems include both rift and convergent margin environments. since these different settings evolved through different thermal histories, it is likely that the possible thermal conditions that may have affected the faroe island basalts, will be represented by at least some of the data recorded in the published studies. the range of fluid compositions at the various sites is broad. the cerro prieto fluids were concentrated solutions with high total dissolved solids and salinities, while many of the solutions reported from the new zealand region, particularly within the broadlands-ohaki (hedenquist 1990) and wairakei areas, included co2-rich and neutral-ph chloride waters and co2-poorer fluids occurred within the iceland system. thus, the published regeus bulletin no 9 7 juli.pmd 07-07-2006, 14:19110 111 ports examined include a range of solutions that are likely to encompass those that may have been present during alteration of the faroe island basalts. clear differences exist between sites with regard to the depth and extent of secondary mineral development, reflecting the effects of these combined intensive and extensive variables (i.e. t, bulk composition, fluid composition etc.). by considering this broad range of systems, it is possible to develop some insight into the extent to which differing geothermal and chemical conditions influenced the development of the mineral associations and how that influence is expressed at the lopra-1/1a site. comparison of the lopra-1/1a suite with these reported mineral parageneses should provide a strong bound to the thermal gradient inferred from these data. in this study, attention is focused on the calc–silicate mineral suite, which is comprised of the components cao–na2o–al2o3–sio2–h2o–co2. although potassium may play an important role in some of these mineral phases, particularly in zeolites where it may substitute for na and ca, it was not considered in this study because it is generally low in abundance in minerals that are characteristically part of the calc–silicate series in basaltic systems. the minerals of interest in the calc–silicate system for the purposes of this study are the zeolites, prehnite, calcite and zoisite–clinozoisite (which are proxies in this study for epidote). this system was selected for detailed consideration because it is the most thoroughly characterised for low-grade mineral development. these minerals possess well-characterised structures and compositions. in addition, there has been a long history of research in the geochemical community to derive thermodynamic data for phases in this system (liou 1971; glassley 1974; frey et al. 1991; neuhoff et al. 1997, 2002; fridriksson et al. 2001; neuhoff & bird 2001). although of immense importance in determining relative conditions in shallow (< 3000 m), low temperature (< 150°c) systems, the clay minerals and chlorites exhibit such structural and compositional complexity that the thermodynamic data available for modelling their behaviour remain inadequate. for that reason, they are not considered further in this report, although work continues on them. consideration of the calc–silicate system also eliminates complexities that arise due to the effects of variable oxygen partial pressures, which can dramatically influence the stability of iron-bearing mineral phases. hence, chlorites, smectites, fe–oxy/hydroxides and related phases are not considered here. two exceptions are considered in this paper. pumpellyite, which is noted in several other studies and documented as a mineral phase of limited distribution at lopra-1/1a, is considered here as part of the paragenetic assemblage, but does not play an important role in establishing the conclusions presented later. prehnite is also considered here and does possess limited solid solution with an fe3+ end member. measured mole fractions in a limited suite of analysed prehnites (unpublished data 1999, r. waagstein) average 0.08, with a range from 0.00 to 0.20 for 18 samples. rose & bird (1987) have shown that solid solution of as little as 10% of the fe end member in al-rich prehnite can significantly affect prehnite stability. although the majority of prehnites analysed in the lopra-1/1a rocks fall below this value, the impact of this effect must be borne in mind and is discussed later in this paper. although the stability fields of many of these minerals are reasonably well established for their ideal compositional end-members, each of these minerals belongs to a solid solution series. generally, there are very little or no quantitative data available regarding the actual compositions of mineral phases in the low-grade rocks described in the referenced reports. in addition, thermodynamic mixing properties of the solid solutions are generally not available. hence, when comparing stability relationships from one locality to another, it must be borne in mind that uncertainties of unknown magnitude are inherent in the comparison due to possible differences in the compositions of the minerals. thermodynamic calculations once mineral assemblages and distributions were compiled, the sensitivity of mineral development to thermal conditions and composition of coexisting fluids was modelled. this effort was undertaken because textural and compositional properties of these secondary minerals attest to the importance of mass transport involving carbonate–bicarbonate-bearing aqueous fluids. the thermodynamic properties of such solutions influence strongly the stability fields of the minerals and can thus be an additional means of placing limits on the physical conditions at the time of mineral growth. the calculations employed the aqueous speciation/reaction progress software eq3/6 (wolery & daveler 1992), using the .com database. the modelling was accomplished by performing speciation calculations over a range of temperatures and compiling the affinities of the possible solid phases that may develop in this system. affinity here is defined as: a = 2.303rt log(q/k) where a is the affinity (in calories), r is the universal gas geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19111 112 constant (1987 calories/mole-degree kelvin), t is temperature (kelvin), q is the activity product for the relevant species in the applicable hydrolysis reaction and k is the equilibrium constant for that same reaction. affinities greater than zero identify mineral phases that are supersaturated in the water at the specified conditions and affinities less than zero identify mineral phases that are undersaturated for those same conditions. positive affinities thus correlate with minerals that would be expected to precipitate from solution or form from mineral reactions in the rock, whereas negative affinities indicate that the respective mineral phase will dissolve, if present. particular attention was given to the development of prehnite since its compositional variability is less than that of the zeolites and its thermodynamic properties are better constrained. the affinities were calculated assuming in all cases that the system was saturated in quartz, laumontite and albite, since these phases coexist with prehnite (see below). these solids were used to constrain the activities of aqueous sio2, al3+ and na+, respectively. the same simulations were repeated assuming that calcite was present as a control for ca++ activity to determine the sensitivity of the results to this change in the system constraints. at the beginning of all of the simulations, it was assumed that the hydrogen ion activity was near neutral at the temperature considered. the initial fluid composition (a dilute, neutral-ph water at the temperature considered) was not in equilibrium with the constraining mineral phases but, for each simulation, was allowed to evolve toward equilibrium with the constraining mineral phases. the equilibrium fluid composition that evolved thus represented the composition of an aqueous fluid in equilibrium with the constraining phases and was the beginning point for further simulations that considered the effects of temperature and other compositional variables. the sensitivity of the results to variations in total cl– and hco3 – was also considered. in this case, the simulations were conducted for cl– concentrations between 14 mg/l and 14.410 mg/l, and hco3 – concentrations between 10 mg/l and 1000 mg/l. this range of values was selected because it encompasses the vast majority of water compositions from hydrothermal systems around the world (see compilations and discussions in roedder 1972; ellis & mahon 1977; arnorsson et al. 1983; fournier 1985). results the depth intervals over which individual minerals occur at the lopra-1/1a site are summarised in fig. 1. noteworthy in this compilation is that the progression with depth of the zeolite sequence is consistent with that from other localities (see summaries below under ‘compiled data’), and that epidote does not occur, even at the deepest levels. also of significance is that most of the minerals persist over depth intervals that exceed significantly any other reported occurrence for that mineral. compiled data the published temperature–depth data compiled from iceland (kristmannsdóttir & tomassón 1976), japan (seki et al. 1969; boles 1981), cerro prieto (bird et al. 1984), wairakei (steiner 1977) and toa baja (cho 1991) are shown in figs 2–4. for each location, the depth interval over which a mineral occurs is indicated by connected symbols that link the high and low temperature and depth points that define the extent of the mineral phase. figures 2–4 also show the depth intervals over which mesolite, stilbite, heulandite, laumontite and prehnite occur in the lopra-1/1a samples (jørgensen 1984, 1997). the lopra-1/1a depth–temperature relationships were constrained to be consistent with the following criteria: 0 1000 2000 3000 4000 pr eh ni te h eu la nd ite m or de ni te sc ol ec ite m es ol ite st ilb ite a na lc im e t ho m so ni te w ai ra ki te la um on tit e pu m pe lly ite d ep th (m et re s be lo w s ur fa ce ) bottom of hole fig. 1. summary of depth distributions for minerals reported in the lopra-1/1a samples (compiled from jørgensen 1984, fig. 4; 1997, fig. 1). zero depth corresponds to the ground surface at the drill site. the bottom of the well is indicated. minerals are arranged along the horizontal axis in a sequence of increasing depth to the right. the depth intervals correspond to the reported occurrences where the individual minerals are most abundant. in some instances, spot occurrences of minerals occur outside the indicated intervals. such occurrences can result from local variations in rock or fluid chemical conditions, or the consequences of locally controlled reaction kinetics, and are not plotted here. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19112 113 1. coexistence of analcime and albite is constrained by cho (1991) to temperatures less than c. 120°c. since albite is ubiquitous in the lopra-1/1a volcanics, the maximum depth occurrence for analcime (c. 1850 m) is assumed to mark the c. 120°c isotherm. 2. laumontite coexisting with prehnite is constrained to temperatures less than 160°c (varna 1989). since laumontite and prehnite occur together over a depth of more than 1000 m and extend to the bottom of the lopra-1/1a hole, this constraint would place the base of the studied sequence at temperatures less than 160°c. 3. epidote is considered to require minimum temperatures for development of 200°c (bird et al. 1984). the exception to this would be systems rich in fe3+ (varna 1989), which the lopra-1/1a basalts are not. epidote is not reported within the lopra-1/1a rocks, hence the bottom-hole temperature must be less than 200°c. 4. pumpellyite requires temperatures in excess of 125°c for stable growth (evarts & schiffman 1983; bevins fig. 2. temperature–depth distributions reported from active thermal systems for the zeolites chabazite, scolecite–mesolite, mordenite, stilbite and heulandite. lines between points indicate the temperature–depth intervals over which the minerals are reported to occur. data sources are: kristmannsdóttir & tomassón 1976 for iceland; seki et al. 1969 and boles 1981 for japan; bird et al. 1984 for cerro prieto, baja california; steiner 1977 for wairakei, new zealand; cho 1991 for toa baja, puerto rico. the solid line labelled lopra is the geothermal gradient derived in fig. 2, with the depth intervals for lopra mesolite, stilbite and heulandite indicated. thor., thorlakshofn, iceland; reyk., reykjavik, iceland; nesj., nesjavellir, iceland. fig. 3. temperature–depth distribution for laumontite and prehnite. laumontite occurrences are from iceland, japan and toa baja, and prehnite from iceland, toa baja and cerro prieto (see fig. 3 for references and abbreviations). also shown for comparison is the inferred temperature–depth distribution for the same lopra minerals along the derived geothermal gradient (fig. 2). prehnite epidote depth (metres below surface) te m pe ra tu re (° c ) reyk. toa baja nesj. cerro prieto krafla thor. wairakei 4000 200 100 400 300 0 lo pra 0 1000 30002000 0 50 100 150 200 250 0 1000 2000 3000 4000 depth (metres below surface) te m pe ra tu re (° c ) low t limit of epidote high t limit of analcime + albite low t limit of pumpellyite et al. 1991). the first appearance of pumpellyite is at a depth of c. 2300 m, thus constraining the 125°c isotherm to be near this depth. these observations were used to construct a palaeogeotherm (fig. 3). in developing this palaeogeotherm, points 1 (the constraint on analcime and albite coexistence) and 4 (the minimum temperature for pumpellyite development) were accepted without qualification. it was also assumed that the mean annual surface temperature was 10°c and that the bottom-hole temperature was c. 200°c. the 200°c bottom-hole temperature, which exceeds the 160°c constraint inferred from coexistence of prehnite and laumontite (point 2), was used to assure a conservative estimate of maximum thermal conditions and represents a compromise between points 2 and 3. in other words, the temperature gradient developed by this approach will overestimate maximum likely thermal conditions. the resulting geothermal gradient is linear. least squares regression of the data points gives a correlation of fit of 0.9949 and a gradient of 0.05°c/m, or 50°c/km. using this geothermal gradient, the depth intervals for mesolite, stilbite and heulandite were plotted to be consistent with the permissible measured distance over which these minerals occur. laumontite and prehnite were placed to be consistent with the implied thermal gradient and temperature constraints, as described above. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19113 114 250 200 150 100 50 0 depth (metres below surface) chabazite mordenite stilbite heulandite scolecite-mesolitereyk. toa baja thor. nesj. krafla thor. lo pr a te m pe ra tu re ( °c ) 0 1000 2000 3000 4000 this reconstruction provides a conservative estimate of the temperature gradient only if the extent of surface erosion since mineral development is small and if there has been minimal tectonic rotation of the volcanic sequence. the consequence of these points is elaborated on below. the following observations are significant for reconstructing conditions recorded in the lopra-1/1a samples. 1. the zeolite group of minerals is stable at temperatures throughout the range 40°c to 210°c (figs 2–4). the only reported occurrence of zeolites at higher temperatures is from the wairakei, new zealand, geothermal field, where wairakite is stable at temperatures of 240°c to 250°c, where it coexists with epidote. this is not an assemblage reported from lopra-1/1a. the corresponding geothermal gradients for wairakei range from a high of > 400°c/km to a low of c. 40°c/km. the highest temperature gradients require active volcanic/ magma systems and are not typical of most environments. nevertheless, the stability relationships for minerals from these systems provide useful information for defining thermal stability limits for the minerals being considered. it should be noted, too, that the higher temperature conditions likely reflect convective hydrothermal environments with highly non-linear geothermal gradients. inevitably, lower geothermal gradients result in a particular mineral being observed over a much longer interval. this then implies that, for a given combination of rockand fluid-compositional characteristics, the lower the temperature gradient, the greater will be the depth range of a borehole over which a particular mineral will occur. 2. although local conditions (such as rock composition, coexisting fluid chemistry, local gas chemistry) at each site determine the exact zeolite sequence, the sequence of minerals generally follows one in which zeolites with high contents of molecular water (e.g. chabazite, scolecite, mesolite) are progressively replaced by zeolites with lower contents of molecular water (e.g. heulandite and laumontite) at higher temperatures. 3. in all cases considered, the assemblage prehnite–laumontite formed near the upper stability field of the zeolites and prior to the appearance of epidote. the temperature range for stable laumontite is in the range 70°c to 200°c. as noted by surdam (1973) and bird et al. (1984), prehnite–laumontite relationships are sensitive to the activity ratio [ca++]/[h+]2 in the fluid fig. 5. temperature constraints for the indicated mineral associations or occurrences. see text for sources and assumptions. the straight line is a least squares fit to the data points. the uncertainty bars for the analcime + albite ‘out’ and the pumpellyite ‘in’ data points span 25°c, and are presented only as an inferred, reasonable uncertainty envelope, in the absence of any available analytical data. the bar associated with the epidote lower t limit indicates the range of possible bottom hole metamorphic temperatures, based on the alternative constraint that the maximum temperature for laumontite coexisting with prehnite is 160°c. see text for further details. fig. 4. temperature–depth distribution for prehnite and epidote. epidote occurrences are from iceland, cerro prieto, toa baja and wairakei (see fig. 3 for references). also shown is the inferred temperature–depth distribution for lopra prehnite along the derived geothermal gradient (fig. 2). prehnite laumontitereyk. toa baja cerro prieto nesj. krafla thor. wairakei japan lo pra depth (metres below surface) te m pe ra tu re (° c ) 4000 200 100 400 300 0 0 1000 30002000 geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19114 115 phase. variation in fluid chemistry is thus the likely cause for the broad temperature interval observed for laumontite stability. 4. in all cases, prehnite first forms at lower temperatures than epidote. however, both occur within the higher temperature range of the zeolites and are stable beyond the zeolite field (fig. 4). prehnite, for example, is reported to be stable in the temperature range 125°c to 340°c in the reports referenced in this study. this temperature interval is the same as reported for the stable presence of epidote, although the lower temperature occurrences of epidote are in systems that have high ca++ and fe3+ activity. the mineral sequence recorded in the lopra-1/1a well (fig. 1) is typical of that reported in other geothermal systems. the zeolite sequence follows the pattern of generally decreasing molecular water content with increasing depth, reflecting the impact of elevated temperatures at deeper levels in the borehole. this observation is generally consistent with the view that the thermal history experienced by these basalts was relatively simple. the highest temperature mineral assemblage that has developed is the prehnite–laumontite assemblage that is reported from the depth interval 2100 m to 3500 m. this assemblage clearly must extend beyond the bottom of the hole to an unknown depth. nevertheless, the 1400 m length of this assemblage is one of the longest such intervals reported anywhere in the world. by comparison, the toa baja prehnite–laumontite zone, the longest interval reported for these minerals, has a total length of about 850 m, and a geothermal gradient of between 50°c/km and 70°c/ km. the inferred temperature interval over which the prehnite–laumontite association formed at lopra-1/1a is inferred to be approximately 120°c to 200°c. epidote does not occur in any of the samples from the lopra-1/1a suite. figure 4 shows that this would require the bottom hole temperature not to exceed c. 250°c to 350°c, which appears to be the temperature interval over which epidote is consistently observed, although lower temperature occurrences have been reported, for example at thorlakshofn and reykjavik in iceland and at toa baja. as noted above, it is inferred that epidote will not form at temperatures less than c. 200°c under conditions of low to moderate fe3+ and ca++ activity. it is thus assumed that the iceland and toa baja occurrences reflect chemical environments that satisfy these conditions. the vestmanna-1 hole, which was also part of the drilling programme (jørgensen 1984, 1997) contains mineral assemblages typical of the shallowest levels of hydrothermal systems and overlap those of the lopra-1/1a sequence. if these mineral assemblages developed simultaneously, the computed geothermal gradient for the lopra-1/1a sequence would have to be considered a maximum. however, uncertainty exists regarding whether these mineral sequences for these two drill holes are coeval. thermodynamic calculations a suite of thermodynamic calculations, using the code eq3/6, was completed to determine the chemical conditions in the fluid phase that would constrain development of the mineral assemblage prehnite–laumontite–quartz– albite–calcite found in the wells. in these calculations, it was assumed that sodium, aluminium, calcium and silica aqueous concentrations are constrained by equilibrium with albite, laumontite, calcite and quartz, respectively. the calculated saturation state of the solution with respect to prehnite was monitored, as temperature and bicarbonate and chloride concentrations were changed. co2 partial pressure was allowed to evolve in response to the equilibrium conditions and monitored to assure that it remained within ‘real world’ bounds. by noting the temperature fig. 6. calculated lower thermal stability limit of prehnite coexisting with albite–calcite–quartz–laumontite, as a function of hco3 – and cl– concentrations in the coexisting aqueous phase. contours on the stability limit surface are labelled in degrees centigrade. the mineral assemblage albite–calcite–quartz–laumontite was used in the calculations because it represents the highest temperature mineral assemblage observed in the bottom of the lopra-1/1a hole. 120°c100°c 160°c 140°c 180°c 200°c 220°c cl–(mg/l) 0 1.0 0.8 0.6 0.4 0.2 0 h c o – (m g/ l) 3 3000 15 00012 00090006000 geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19115 116 and bicarbonate and chloride concentrations at which the solution became saturated in prehnite, it is possible to delineate those conditions that bound the stability field for the prehnite-bearing mineral assemblage. the results of the calculations are presented in fig. 6, which shows the contoured temperature surface for the stability of prehnite coexisting with a bicarbonate–chloride solution in equilibrium with laumontite–calcite– quartz–albite. the contours map the minimum temperature required for prehnite stability in this system. it must be emphasised that the exact location of these contours is somewhat imprecisely known, due to uncertainty in the thermodynamic data. the uncertainty in the bicarbonate values is approximately ± 50 mg/l, based on interpolations between simulations. these results show that prehnite stability is only slightly sensitive to the solution salinity (as indicated by the effect of variation in the chloride ion, cl–), but is very sensitive to the solution carbonate/ bicarbonate concentration. this behaviour reflects the strong coupling between these variables and ca speciation and ph. the more concentrated the solution in terms of carbonate/bicarbonate, the higher the temperature necessary to achieve prehnite stability. these calculations suggest that the wide range of prehnite thermal stability observed in natural systems (figs 2, 5) is due, at least in part, to differences in fluid composition from one location to another. this probably is true for other minerals in this calc–silicate suite as well. as documented by rose & bird (1987), the redox state and iron content of the fluid will also be an important variable in controlling prehnite stability, due to the effect of fe3+ substitution for al in the prehnite structure. salinities determined from a preliminary fluid inclusion study of the lopra-1/1a samples (konnerup-madsen 1998) gave cl– concentrations of between 0.167 and 1.49 equivalent weight per cent nacl, which is approximately 1000 to 9000 mg/l cl–. the analytical bicarbonate ion concentrations with this salinity in natural solutions in hydrothermal systems and at these temperatures and pressures are usually in the range of 200 to 800 mg/l (see compilations and discussions in roedder 1972; ellis & mahon 1977; arnorsson et al. 1983; fournier 1985) although the actual hco3 – concentrations in the reservoirs will be lower than this value and be controlled by co2 fugacity. this implies (fig. 6) that the mineral association prehnite–laumontite–calcite–quartz formed at temperatures within the range of approximately 140°c to 205°c. this temperature interval is contained within the range of prehnite stability noted in other hydrothermal systems (see figs 2, 5) and is thus consistent with natural occurrences of this assemblage. it is also broadly consistent with the inference from phase relationships described above, in which it is suggested that this assemblage spans the temperature interval of approximately 120°c to 200°c. discussion and conclusions secondary mineral assemblages documented for the basalts recovered from the lopra-1/1a well are similar to those reported from other hydrothermal systems. both the specific mineral occurrences and the relative sequence of mineral stabilities define a systematic distribution that records increasing temperature with depth. the absolute length of individual mineral zones, however, is greater than at other well-documented sites, and suggests that the geothermal gradient at the time of mineral development was low. the mineral associations, complemented by thermodynamic calculations of fluid-rock equilibrium relationships, suggest that the temperature at the bottom of the well did not exceed 200°c, implying a maximum thermal gradient of 50°c/km (assuming a surface temperature in the range of 10 to 25°c). this gradient was constructed based on the assumption that the mineral zones are approximately horizontal. there is currently no structural data available to suggest this assumption is far from accurate, but it remains to be established conclusively. furthermore, it is also assumed that the total stratigraphic thickness at the time of mineral development did not greatly exceed that exposed and inferred today. this assumption is reasonable, based on the correlations established by larsen et al. (1999) between the east greenland volcanic complex and the faroe islands. the correlations indicate that the current thickness of basalts in the faroe islands is probably close to that which was originally erupted. it has previously been suggested that mineral development may have occurred in several discrete episodes (jørgensen 1984, 1997). such an interpretation makes more complex the sequence and timing of mineral growth and may change the absolute depth intervals over which specific mineral associations formed within a given time period. this, in turn, would require reconsideration of the temperature history since such an observation could result only in shorter absolute depth intervals for each mineral development period. in this scenario, the currently observed distribution of minerals would represent the sum of the depth intervals over which an individual mineral formed at different time periods, assuming that no single episode of mineral development obliterated evidence of previous distributions of secondary mineral development. nevertheless, the conclusion that the bottom hole temgeus bulletin no 9 7 juli.pmd 07-07-2006, 14:19116 117 perature did not exceed 200°c would still be valid, since that is based on the mineral association calcite–laumontite–prehnite–quartz, the temperature limit of which is constrained by laumontite and prehnite thermal stability and fluid composition effects. comparison of the derived geothermal gradient in the faroes with those reported for the atlantic margin region north of the united kingdom and in east greenland demonstrates a striking consistency that constrains evolution of the geothermal history in this region. green et al. (1999) used fission track data from apatites as well as vitrinite reflectance data from a series of wells in the eastern north atlantic province to determine palaeogeothermal gradients. they reported geothermal gradients of between 35°c/km and 90°c/km, with the vast majority of the region falling within the lower portion of the range. neuhoff et al. (1997) concluded that the zeolite facies metamorphism that affected east greenland flood basalts during initial opening of the northern north atlantic resulted from recrystallisation associated with a geothermal gradient of 40 ± 5°c/ km. the regional heat flow they derived from this conclusion is consistent with that reported from a study of metamorphic recrystallisation (manning et al. 1993). all of these values effectively bracket the inferred geothermal gradient in the faroe islands and argue for early development of relatively low geothermal gradients that persisted for some time in these regions. these results, and those of larsen et al. (1999), provide conceptual constraints on models of the thermal evolution of this part of the northern north atlantic province during early continental separation and basin development and argue for regions of low geothermal gradients that were not overprinted by later high heat-flow periods. as a word of caution, it should be noted that these conclusions are based on the simplifying assumption that linear geothermal gradients existed during mineral growth in this region. there is substantial evidence in geothermal systems, however, that complex geothermal gradients commonly develop, such that temperature reversals or near isothermal conditions may develop in response to the local thermal–hydrological regime, particularly in environments dominated by convection-driven fluid flow. although such features usually develop in regions of high heat flow and are not characteristic of environments such as the faroe islands region where heat flow is inferred to be low, evidence is currently inadequate to rule out this possibility conclusively. to evaluate the extent to which such behaviour occurred in the faroe islands volcanic province, a more detailed examination of mineral composition characteristics and distributions would be required, coupled with a more detailed modelling effort. acknowledgements regin waagstein kindly provided timely access to thin sections, mineral composition data and mineral distribution data, as well as informative discussions. his assistance greatly aided this effort. extensive comments from dennis bird and bruce christenson led to significant improvements in earlier versions of the manuscript, and are gratefully acknowledged. the editorial wisdom of james a. chalmers significantly improved the presentation and style of this paper. references arnorsson, s., gunnlaugsson, e. & svavarsson, h. 1983: the chemistry of geothermal waters in iceland. ii. mineral equilibria and independent variables controlling water compositions. geochimica et cosmochimica acta 47, 547–566. bevins, r.e., rowbotham, g. & robinson d. 1991: zeolite to prehnite–pumpellyite facies metamorphism of the late proterozoic zig-zag dal basalt formation, eastern north greenland. lithos 27, 155–165. bird, d., schiffman, p., elders, w.a., williams, a.e. & mcdowell, s.d. 1984: calc–silicate mineralization in active geothermal systems. economic geology 79, 671–695. boles, j.r. 1981: zeolites in low grade metamorphic rocks. in: mumpton, f.a. (ed.): mineralogy and geology of zeolites. mineralogical society of america reviews in mineralogy 4, 103– 135. cho, m. 1991: zeolite to prehnite–pumpellyite facies metamorphism in the toa baja drill hole, puerto rico. geophysical research letters 18, 525–528. ellis, a.j. & mahon, w.a.j. 1977: chemistry and geothermal systems, 392 pp. new york: academic press. evarts, r.c. & schiffman, p. 1983: submarine hydrothermal metamorphism of the del puerto ophiolite, california. american journal of science 283, 289–340. fournier, r.o. 1985: continental scientific drilling to investigate brine evolution and fluid circulation in active hydrothermal systems. in: raleigh, c.b. (ed.): observation of the continental crust through drilling i, 98–122. berlin: springer-verlag. frey, m., de capitani, c. & liou, j.g. 1991: a new petrogenetic grid for low-grade metabasites. journal of metamorphic geology 9, 497–509. fridriksson t., neuhoof, p.s., arnorsson, s. & bird, d.k. 2001: geological constraints on the thermodynamic properties of the stilbite–stellerite solid solution in low-grade metabasalts. geochimica et cosmochimica acta 65, 3993–4008. glassley, w. 1974: a model for phase equilibria in the prehnite– pumpellyite facies. contributions to mineralogy and petrology 43, 317–332. green, p.f., duddy, i.r., hegarty, k.a. & bray, r.j. 1999: early geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19117 118 tertiary heat flow along the uk atlantic margin and adjacent areas. in: fleet, a.j. & boldy, s.a.r. (eds): petroleum geology of northwest europe, proceedings of the 5th conference, 349– 357. london: geological society. hedenquist, j.w. 1990: the thermal and geochemical structure of the broadlands–ohaaki geothermal system, new zealand. geothermics 19, 151–185. jørgensen, o. 1984: zeolite zones in the basaltic lavas of the faeroe islands. in: berthelsen, o., noe-nygaard, a. & rasmussen, j. (eds): the deep drilling project 1980–1981 in the faeroe islands. annales societatis scientiarum faroensis. supplementum 9, 71–91. jørgensen, o. 1997: zeolites and other secondary minerals in cavities and veins, lopra-1/1a well, faroe islands, 1996, 8 pp. + plates. unpublished report, technical studies prepared for dansk olie og gasproduktion a/s, copenhagen, denmark (in archives of geological survey of denmark and greenland, geus report file 26129). konnerup-madsen, j. 1998: a preliminary examination of fluid inclusions in vug and fracture-filling quartz and calcite from lopra-1/1a, faroe islands, 5 pp. unpublished report, geological survey of denmark and greenland, copenhagen. kristmannsdóttir, h. & tomassón, j. 1976: zeolite zones in geothermal areas in iceland. in: sand, l.b. & mumpton, f.a. (eds): natural zeolites; occurrence, properties, use, 277–284. oxford: pergamon press. larsen, l.m., waagstein, r., pedersen, a.k. & storey, m. 1999: trans-atlantic correlation of the palaeogene volcanic successions in the faeroe islands and east greenland. journal of the geological society (london) 156, 1081–1095. liou, j.g. 1971: synthesis and stability relations of prehnite, ca2al2si3o10(oh)2. american mineralogist 56, 507–531. manning, c.e., ingebritsen, s.e. & bird, d.k. 1993: missing mineral zones in contact metamorphosed basalts. american journal of science 293, 894–938. neuhoff, p.s. & bird, d.k. 2001: partial dehydration of laumontite; thermodynamic constraints and petrogenetic implications. mineralogical magazine 65, 59–70. neuhoff, p.s., watt, w.s., bird, d.k. & pedersen, a.k. 1997: timing and structural relations of regional zeolite zones in basalts of the east greenland continental margin. geology 25, 803–806. neuhoff, p.s., kroeker, s., du, l.s., fridriksson, t. & stebbins, j.f. 2002: order/disorder in natrolite group zeolites: a 29si and 27al mas nmr study. american mineralogist 87, 1307–1320. neuhoff, p.s., stebbins, j.f. & bird, d.k. 2003: si-al disorder and solid solutions in analcime, chabazite, and wairakite. american mineralogist 88, 410–423. neuhoff, p.s., hovis, g.l., balassone, g. & stebbins, j.f. 2004: thermodynamic properties of analcime solid solutions. american journal of science 304, 21–66. roedder, e. 1972: composition of fluid inclusions. in: data of geochemistry. u.s. geological survey professional paper 400jj, 164 pp. rose, n.m. & bird, d.k. 1987: prehnite-epidote phase relations in the nordre aputiteq and kruuse fjord layered gabbros, east greenland. journal of petrology 28, 1193–1218. seki, y., onuki, h., okumura, k. & takashima, i. 1969: zeolite distribution in the katayama geothermal area of japan. japanese journal of geology and geography 40, 63–79. steiner, a. 1977: the wairakei geothermal area, north island, new zealand: its subsurface geology and hydrothermal rock alteration. new zealand geological survey bulletin 90, 136 pp. surdam, r.c. 1973: low-grade metamorphism of tuffaceous rocks in the karmutsen group, vancouver island, british columbia. geological society of america bulletin 84, 1911–1922. varna, c.l. 1989: mineral reactions and controls on zeolite-facies alteration in sandstones of the central transantarctic mountains, antarctica. journal of sedimentary petrology 59, 688– 703. waagstein, r., guise, p. & rex, d. 2002: k/ar and 39ar/40ar whole-rock dating of zeolite facies metamorphosed flood basalts: the upper paleocene basalts of the faroe islands. in: jolley, d.w. & bell, b.r. (eds): the north atlantic igneous province: stratigraphy, tectonic, volcanic and magmatic processes. geological society special publication (london) 197, 219–252. wolery, t.j. & daveler, s.a. 1992: eq6, a computer program for reaction path modeling of aqueous geochemical systems: theoretical manual, user’s guide, and related documentation. lawrence livermore national laboratory ucrl-ma-110662 part iv, 338 pp. manuscipt received 22 december 1999; revision accepted 26 may 2005. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19118 geological survey of denmark and greenland bulletin 42, 2018, 65-84 65 diagenesis of upper jurassic sandstones of the blokelv-1 core in the jameson land basin, east greenland mette olivarius, rikke weibel, niels h. schovsbo, dan olsen and claus kjøller petrographic analysis combined with x-ray diffraction are used to identify the diagenetic changes that have affected the porosity and permeability of gravity-flow sandstones of the oxfordian–volgian hareelv formation in the cored blokelv-1 borehole in jameson land. kaolinite replacement of albite grains probably occurred early after deposition and microquartz coatings formed under shallow burial. at deeper burial, illite and quartz formed from kaolinite and k-feldspar. pervasive ankerite cement formed in the finest grained sandstones and may have formed at the expense of early calcite cement. quartz overgrowths are volumetrically small, partly due to inhibition by microquartz coatings and partly due to limited residence time during deep burial. the succession reached the maximum burial depth of c. 2.8 km during the late eocene. basaltic material was intruded into the sediments during the early eocene and the enhanced heat flow accelerated diagenesis in the close vicinity of the intrusions, which have thicknesses of up to 2 m. most of the sandstones have porosities between 14.4 and 25.7% and permeabilities between 0.4 and 411.9 md; this variation resulted from a combination of microquartz coatings and clay minerals. however, the intrusion-influenced sandstones and the ankerite-cemented sandstones have lower porosity and permeability. keywords: east greenland, hareelv formation, katedralen and sjællandselv members, oxfordian–kimmeridgian, petrography, mineralogy, reservoir quality. ___________________________________________________________________________ geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: mol@geus.dk the blokelv-1 fully cored borehole was drilled with almost full recovery to a total depth of 234 m in southern jameson land, east greenland within the central part of the jameson land basin (fig. 1; bjerager et al. 2018a, this volume). laminated mudstones dominate the cored succession of jurassic sediments referred to the hareelv formation. sandstone layers make up 36% of the core and are generally 1–7 m thick, and heterolithic bedded intervals make up 11% of the cored succession (fig. 2; bjerager et al. 2018a, this volume). good reservoir quality is present in most of the sandstones, and the interbedded mudstones constitute rich gas/oil-prone source rocks (boje sen-koefoed et al. 2018, this volume). thus, the blokelv-1 succession may act as an analogue for similar sandstone– mudstone complexes in the north atlantic region. the purposes of this study are: (1) to determine the diagenetic development of the sandstones of the hareelv formation in order to explain their large variation in porosity and permeability; (2) to relate the observed mineralogical changes to diagenetic processes caused by burial and igneous intrusion; and (3) to interpret their effect on porosity evolution in the sandstones. geological setting thick post-caledonian sedimentary successions were deposited in east greenland in response to basin development during thermal subsidence and rifting (stemmerik et al. 1992; surlyk 2003). the devonian–jurassic © geus, 2018. geological survey of denmark and greenland bulletin 42, 65–84. available at: www.geus.dk/bulletin42 mailto:mol@geus.dk http://www.geus.dk/bulletin42 6666 succession in the jameson land basin has a thickness of up to about 18 km (larsen & marcussen 1992), of which the jurassic sediments comprise 2 km (bjerager et al. 2018a, this volume). younger sediments were largely eroded during cenozoic uplift of the jameson land area (mathiesen et al. 2000; hansen et al. 2001), although the cretaceous hesteelv formation is preserved as a thin cover in southernmost jameson land (fig. 1). the exhumation of the studied succession began in the late eocene and renewed uplift took place from the late miocene and onwards (green & japsen 2018, this volume); these are widely known events along the east greenland margin ( japsen et al. 2014). organic-rich mudstones were deposited in central areas of the jameson land basin during the late jurassic; anoxic conditions prevailed in relatively deep water, well below storm wave base. such mudstones comprise the majority of the hareelv formation in the blokelv-1 core (fig. 2; surlyk et al. 2007; bjerager et al. 2018a, this volume). the interbedded sandstones were deposited from gravity flows in lower slope and basin-floor settings (fig. 3). most of the sandstone beds in the blokelv-1 core belong to the oxfordian–kimmeridgian katedralen member of the hareelv formation; some of these gravity-flow sands were remobilised and injected into the surrounding mudstones. the coarsest and thickest gravity-flow sandfig. 1. geological map of jameson land in east greenland showing the location of the fully-cored blokelv-1 borehole. the triassic–cretaceous sediments are intersected by numerous igneous intrusions. based on the digital greenland geological map at a scale of 1:500 000 (data.geus.dk/ map2/geogreen) and the printed map series at a scale of 1:100 000. ; ; ; ; ; ; 70°30'n 71°n olympen fm fossilbjerget fm / pelion fm neill klinter gp kap stewart gp triassic basement hareelv fm, katedralen mb major dyke/sill ice quaternary hesteelv fm raukelv fm hareelv fm, salix dal mb hareelv fm, sjællandselv mb ?devonian blokelv-1 20 km 23°w24°w jameson land h urry in le t mol dia 1 s c o r e s b y s u n d http://data.geus.dk/map2/geogreen http://data.geus.dk/map2/geogreen 67 stone bed in the blokelv-1 core is present in the uppermost part where it is referred to the volgian sjællandselv member of the hareelv formation (fig. 2). the sediment in the gravity-flow sandstones of the katedralen member is interpreted to have been fed by collapse of drowned shelf-edge deposits exemplified by the zeus member of the oxfordian olympen formation situated at the northern margin of the basin (fig. 3a; surlyk et al. 2007; bjerager et al. 2018b, this volume). the latter sand unit represents a southwards-prograding delta that accumulated under lowstand conditions prior to deposition of the katedralen member (larsen & surlyk 2003). the katedralen member accumulated during a major transgressive pulse that culminated in the late kimmeridgian. the subsequent relative sea-level fall in the early volgian was heralded by the influx of gravityflow sands forming the sjællandselv member, derived from shelf-edge sands analogous to the younger raukelv formation (fig. 3b) that prograded east and southwards into the basin (surlyk 2003; bjerager et al. 2018b, this volume). the provenance of the hareelv formation comprises archaean–palaeoproterozoic crystalline complexes, meso–neoproterozoic metamorphic rocks and caledonian migmatites and granites; some or all of the material may have been through several cycles of sedimentation prior to deposition in the late jurassic in the jameson land basin (olivarius et al. 2018, this volume). samples and methods the sampled sandstones are grey, moderately to wellsorted and very fineto medium-grained (fig. 4; bjerager et al. 2018a, this volume). analyses of both bulk and clay mineralogy were carried out by x-ray diffraction (xrd) on 22 samples comprising 18 sandstones, one mudstonedominated heterolith and three mudstones. the sandstone intervals were sampled to reflect the full spectrum of facies present in the succession; sandstone samples were also taken close to the contacts with mudstones and igneous intrusions to investigate if the proximity to these had an effect on the diagenesis. thin sections were made of 25 sandstone samples that were selected on the basis of both representivity and diversity. quantitative determination of detrital and authigenic mineral content was made by point counting of thin sections prepared from nine sandstone samples to constrain the mineralogy further, and the selected intervals correspond to nine of the xrd analyses. scanning electron microscopy (sem) analyses were carried out to determine the morphology of the minerals and the relative timing of the diagenetic processes. geochemical analyses of 42 samples are reported by bjerager et al. (2018a, this volume) and the results are used here as input parameters in a principal component analysis (pca) in order to determine the relationships between the geochemistry and the mineralogy. he-porosity, air permeability and grain density measurements are also reported by bjerager et al. (2018a, this volume) and their results from 18 sandstone samples are included here and used in the interpretation of reservoir quality. core scanning data (spectral gamma-ray, bulk density) are also described by bjerager et al. (2018a, this volume) along with the log-derived porosity and permeability logs; three selected intervals of these logs are presented here at high resolution to illustrate the relationships to the mineralogical data. x-ray diffraction bulk mineralogy was measured by xrd on half-core samples of 3–4 cm length. the rim of the core was removed prior to crushing to <63 µm. the bragg–brentano method was applied utilising a bruker advance d8 diffractometer with a lynx-eye detector. the bulk mineralogy was quantified by the rietveld method (rietveld 1969; mccusker et al. 1999). the clay mineralogy was measured by xrd on halfcore samples of 5–6 cm length when sand was dominant and of 1–2 cm length when clay dominated. the samples were gently hand-ground to pass a 250-µm sieve. organic matter was removed in the chemical pre-treatment using naocl at ph 9.0. the samples were dispersed ultrasonically in distilled water to obtain the clay fraction <2 µm for analysis. the fraction >30 µm was removed by sedimentation and the 2–30 µm fraction was separated in a centrifugal particle-size analyser (slater & cohen 1962). the resulting suspension was flocculated in 1 m nacl. excess salt was removed by centrifugation and washing with water and ethanol. the clay fraction was then airdried. three oriented specimens were prepared for each sample by the pipette method as follows: mg-saturated air-dry, mg-saturated with glycerol added to the suspension and k-saturated air-dry heated for 1 hour at 300°c. the xrd analysis was carried out on randomly oriented specimens using a philips 1050 goniometer with fixed divergence, anti-scatter slits and co-kα radiation (pulsehigh selection and fe-filter). an x-ray diffractogram was produced for each of the three saturated specimens for each sample. the discrete minerals were identified 6868 chronostratigraphy lithostrat. m id dl e up pe r lo w er up pe r vo lgi an o xf or di an ki m m er id gia n u pp er ju ra ss ic h ar ee lv fo rm at io n ka te dr ale n m em be r sjæ lla nd se lv m b lo w er 10 depth (m) 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 210 220 230 233.8 be be mudstone heterolith (mudstone/sandstone) sandstone sandstone, remobilised (intruded) lithology igneous intrusion be ripple cross-lamination wavy bedding slump sandstone intrusion small mudstone intraclasts bentonite large mudstone clast coalified wood belemnite ammonite bivalve brachiopod structures, biota parallel lamination/bedding diffuse stratification clay si sand pebbl. mol dia 2 fig. 2. sedimentological log of the blokelv-1 core including chronoand lithostratigraphy. the log is from bjerager et al. (2018a, this volume) and the biostratigraphy is from alsen & piasecki (2018, this volume). 69 from peak positions on the xrd diffractograms (hillier 2000) and the amounts were semi-quantitatively determined as major and minor phases. petrographic analysis polished thin sections were prepared from sandstone samples using blue epoxy for the impregnation to help the identification of open pore space. half of each thin section was etched and stained with sodium cobaltinitrite for k-feldspar identification. quantitative point counting was made by identifying 300 mineral grains in each thin section by the use of a petrographic microscope with polarisation filters. the open pore spaces were also counted. sem analyses of sandstones were made on a philips xl40 sem equipped with thermo nanotrace and pioneer voyager energy dispersive x-ray (edx) detectors. edx was used for making elemental analyses of the grains to give a semi-quantitative identification of the minerals. the sem analyses were performed on carboncoated thin sections and gold-coated rock chips placed on carbon tape. principal component analysis multivariate data analysis was used to classify the lithology types according to their elemental composition by applying principal component analysis (pca). the pca transforms a matrix of measured data x (n samples, p variables) into sets of projection sub-spaces. these are delineated by principal components that are each a linear combination of all p variables and display variance-maximised interrelationships between samples and variables, respectively (esbensen 2010). pca score plots display groupings between samples based on compositional n deep-water marine sandstone bodies drowned (relict) sandy shelf (olympen fm, zeus mb)shallow marine sandstone hareelv fm, katedralen mb raukelv fm hareelv fm, sjællandselv mb offshore marine mudstone inferred coastline transition zone marine sandstone/mudstone heterolith 50 kma b mol dia 3 jameson land traill ø milne land fig. 3. palaeogeographic reconstructions for the katedralen member (a) and sjællandselv member (b) of the hareelv formation (modified after surlyk 2003; see bjerager et al. 2018b, this volume). the facies distribution is tentative in most areas since the distribution in the presentday offshore area is inferred and the succession has been removed by erosion in some of the present-day onshore areas. 7070 similarities, as described by the variable correlations that are shown in accompanying loading plots. furthermore, the proportion of total data-set variance that can be modelled by each component is quantified. all data analyses in this study are based on auto-scaled data [x-x(avr)/ std]. the elements w, co and ta are excluded from the pca since the values are misleading due to contamination from the tungsten carbide mortar. this contamination is rock-type specific since it was greatest for the silica-rich sandstones as they are harder than ankeritecemented sandstones and mudstones. results overall, there is good correlation between the mineralogy measured by bulk xrd and that obtained by point counting of thin sections (figs 5, 6), but it should be noted that the xrd results are in wt% and the petrographic results are in vol%. however, the amount of quartz is always higher in the xrd results than in the petrographic results (8% higher on average), whereas the amounts of clays, micas and feldspars are generally lowest in the xrd results. the three xrd analyses of bulk and clay fraction from 223.5, 221.2 and 218.8 m depth were carried out on samples from the base, middle and top, respectively, of the same turbiditic sandstone layer and show similar mineralogies (figs 2, 5). mineralogical composition the average quartz content of the sandstones is 66 wt% according to the xrd results, whereas quartz comprises 17 wt% of the mudstones (fig. 5). point counting shows that the detrital quartz grains are chiefly monocrystalline. authigenic quartz on average comprises 5% of the sandstones and 9% of the total quartz content (fig. 6), but the content of authigenic quartz is probably underestimated since microquartz could not be identified during point counting. the amount of authigenic quartz is constant up through the core whereas the clay content in the sandstones decreases upwards and is approximately halved from bottom to top. the sum of clay and mica minerals is 14 wt% on average for the sandstones and 69 wt% for the mudstones (fig. 5). the amount of clay minerals usually exceeds the amount of mica minerals, and muscovite is the most common mica mineral (fig. 6). the feldspars amount on average to 9 wt% in the sandstones and mudstones, and albite is more abundant than k-feldspar (fig. 5, table 1). the albite content shows a weak upward increase in the cored sandstones. some of the feldspar that is identified as albite in the xrd rietveld quantification consists of low-ca plagioclase as revealed by sem edx analysis. pyrite amounts on average to 1 wt% in the sandstones and 6 wt% in the mudstones (fig. 5). the edx analyses confirm the xrd results, which show that ankerite is the only carbonate mineral in the sediments. the amount of ankerite is <1 wt% in the mudstones, but large variations in abundance (0–41 wt%) are found in the sandstones (fig. 5). the highest porosity: 17.55% permeability: 30.79 md grain density: 2.65 g/cm3 porosity: 6.37% permeability: 0.05 md grain density: 2.75 g/cm3 mol dia 4 a b 500 µm500 µm fig. 4. typical textures of sandstones from the hareelv formation. a: fine-grained sandstone with open pores indicated by blue epoxy coloration. b: very fine-grained sandstone with pervasive ankerite cementation. in general, high porosities and intermediate to high permeabilities are present in the fineand medium-grained sandstones, whereas the very fine-grained sandstones contain pervasive ankerite cement and thus show poor reservoir quality. 71 contents of organic matter are present in ankerite-cemented sandstones (fig. 6). high ankerite content in the sandstones correlates with high feldspar content and low quartz content (table 1). the average feldspar/quartz weight ratio is 0.09 in sandstones with <5 wt% ankerite, 0.19 in sandstones containing 5–25 wt% ankerite, 0.33 in sandstones with >25 wt% ankerite, and 0.45 in the analysed mudstones. the grain size of the sandstones correlates inversely with the ankerite content so a large amount of ankerite equates to a small grain size and vice versa (table 1). kaolinite is the dominant clay mineral in all lithologies, but it is absent in one sandstone sampled adjacent to an igneous intrusion where illite and mixed-layer clays dominate (fig. 5). mixed-layer clays and illite are present in all lithologies. vermiculite is present in the mudstones and the heterolith that were analysed, but only in two of the sandstones. geochemical composition the first two principal components in the pca model resolve 82% of the total variance in the bulk geochemical data (fig. 7a). the main trend is expressed by the pc1 axis and represents the variation in clay versus silica content, as seen by high positive pc1 loadings for element like al2o3, k2o and most trace elements, and high negative values for sio2. the pc2 axis displays high positive loadings of cao, mgo and mno and high negative loadings of sio2, which reflect the variation in carbonate 7.41 43.16 56.43 56.91 61.96 69.78 74.99 98.92 114.69 116.32 118.72 128.80 139.27 151.71 166.95 178.32 178.48 181.53 218.50 218.80 221.18 223.53 s s s(i) s s s m s h s s s s s s m s s m s s s bulk mineralogy (wt%) clay mineralogy lithology 0 20 40 60 80 100depth (m) clays, micas albite k-feldspar quartz ankerite pyrite kaolinite mixed-layer illite vermiculite bulk mineralogy clay mineralogy sandstone sandstone (near intrusion) heterolith mudstone lithology s(i) m s h x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x x clay quantity major phase minor phase k vim mol dia 5 fig. 5. bulk and clay mineralogy of samples from the blokelv-1 core measured by x-ray diffraction (xrd). the bulk mineralogy was quantified by the rietveld method (rietveld 1969). the clay mineralogy is semi-quantitatively determined as major and minor phases. 7272 versus silica content that is associated with carbonate cementation in some of the sandstones. the different lithologies are clearly separated on the pca score plot (fig. 7b). the sandstones are characterised by negative pc1 score values and the carbonatecemented parts (>5% ankerite) have positive pc2 score values. the mudstones have the highest pc1 scores, reflecting the highest clay and lowest silica content, and the heteroliths have intermediate pca score values compared to sandstone and mudstones, as expected from their intermediate composition. it is evident from the pc2 score values that the mudstones are not cemented with ankerite whereas the heteroliths may be slightly ankerite-cemented (fig. 7b). the large amounts of zr, hf and na2o in the ankerite-cemented sandstones show that they contain more zircon and albite than the uncemented sandstones. correlation between ankerite-cemented sandstones and fe2o3 is not evident due to the high pyrite content of the mudstones. reservoir properties a moderately good correlation (r2 = 0.79) exists between porosity and air permeability in the sandstones of the blokelv-1 core (fig. 8). in general, permeabilities lower than the trendline are related to finer-grained sandstones while permeabilities above the trendline are associated with coarser-grained sandstones. the lowest porosities and air permeabilities of 6.4–9.6% and 0.05–0.21 md, respectively, are found in two ankerite-cemented sandstones which are also characterised by high grain densities of 2.75 g/cm3 (fig. 8). the remaining sandstones without significant ankerite cementation have grain densities of 2.62–2.69 g/cm3, porosities of 14.3–25.7% and air permeabilities of 0.3–411.9 md. the inter-granular volume (igv) is the sum of porosity and authigenic minermudstone <5 wt% 5–25 wt% >25 wt% <1 wt% clay and mica (wt%) 13.7 15.0 16.6 68.6 albite (wt%) 3.8 7.5 7.9 4.9 k-feldspar (wt%) 3.0 3.4 4.4 2.8 quartz (wt%) 76.8 60.1 36.9 17.3 ankerite (wt%) 1.6 13.6 33.4 0.5 pyrite (wt%) 1.1 0.5 0.7 5.9 feldspar/quartz ratio 0.09 0.19 0.33 0.45 grain size (µm) 185 159 105 number of samples 11 4 3 3 sandstone table 1. mineral content and feldspar/quartz ratios versus lithology and ankerite content the values are averaged from bulk xrd analyses. groups by ankerite content mol dia table 1 clay minerals quartz, authigenic ankerite pyrite, framboidal biotite, chlorite, heavy minerals muscovite rock fragments albite, plagioclase k-feldspar quartz, monocrystalline quartz, polycrystalline organic matter 0 20 40 60 80 0 20 40 detrital minerals (%) authigenic minerals (%) 60100 2.0 4.7 0.0 0.0 19.0 0.0 16.3 18.7 1.3 porosity (%) 33.3 21.3 20.7 29.7 33.3 46.0 39.0 40.0 28.3 igv (%) 7.41 61.96 98.92 118.72 128.80 151.71 178.48 181.53 223.53 depth (m) mol dia 6 fig. 6. quantitative bulk mineralogy of sandstones from the blokelv-1 core determined by point counting. the detrital and authigenic minerals are summed to 100%; the sum of porosity and authigenic minerals constitutes the inter-granular volume (igv). 73 als (houseknecht 1987; ehrenberg 1989) and it amounts to 21–46% in the point-counted sandstones (fig. 6). the ankerite-cemented sandstone intervals are clearly recognisable on the bulk-density log due to the high grain density of ankerite (fig. 9). it is evident from the bulk-density log that extensive metre-thick, ankerite-cemented sandstone intervals occur immediately adjacent to sandstones with a low degree of ankerite cementation. the porosity and permeability logs show that porosity and especially permeability are low in the ankeritecemented sandstones. the values are reduced in sandstones above and below cemented intervals as compared to sandstone beds that do not have any intervals with pervasive ankerite cementation and therefore typically have porosities greater than 18% and permeabilities above 10 md (fig. 9). pc1 scores (72%) -10 -5 0 5 10 15 pc 2 sc or es (1 0% ) -4 -2 0 2 4 6 pc1 loadings (72%) -0.2 -0.1 0.0 0.1 pc 2 lo ad in gs (1 0% ) -0.4 -0.2 0.0 0.2 0.4 0.6 0.2 sio2 al2o3, k2o, tio2, cr2o3, tot/s, ree, sc, cs, ga, nb, rb, th, u, v, y, cu, pb, ni, as fe2o3 mgocao na2o p2o5 mno loi ba sr zr mo zn sandstone with >25 wt% ankerite heterolith mudstone sandstone with <5 wt% ankerite sandstone with 5–25 wt% ankerite hf tot/c b a mol dia 7 fig. 7. results of multivariate data analysis (principal component analysis, pca) based on bulk geochemical analyses of the blokelv-1 core (data from bjerager et al. 2018a, this volume). the first and second principal components (pc) are shown. a: on the loading plot, the elements cluster according to their mineralogical association. b: the score plot shows clear separation between each lithology class based on its score value. the sandstones are subdivided by their ankerite content, which was determined by xrd for 18 of the samples. r2 = 0.79 porosity (%) ai r p er m ea bi lit y (m d ) g ra in d en sit y (g /c m 3 ) 2.76 2.72 2.64 2.68 2.60 1000.00 100.00 10.00 1.00 0.10 0.01 sandstone near igneous intrusion a b sandstone with pervasive ankerite cement sandstone with no or patchy ankerite cement 0 30155 10 20 25 0 30155 10 20 25 mol dia 8 fig. 8. he-porosity versus air-permeability (a) and grain density (b) of sandstone samples from the hareelv formation (bjerager et al. 2018a, this volume). the ankerite-cemented sandstones have the lowest porosity and permeability and the highest grain density of all samples. relatively low porosities and permeabilities are also observed in sandstones that have been affected by igneous intrusions, whereas the variation in the remaining sandstones is caused by a combination of microquartz coatings and clay minerals. 7474 150 151 152 153 115 116 117 k2o (%) th (ppm) u (ppm) gamma ray total (api) 166 167 permeability (md) porosity (%) bulk density (g/cm3) 2.2 5 2.5 0 2.7 5 0 1 2 3 0 10 20 0 10 200 10 0 20 0 0. 01 0. 1 1 10 10 0 0 10 20 an ke rit e 41 % an ke rit e 30 % sandstone (ankerite-cemented) mudstone/heterolithsandstone depth (m) core measurement an ke rit e 2% mol dia 9 fig. 9. high-resolution profile of the core spectral gamma ray (grtotal, k, th, u) and bulk density log with the derived porosity and permeability logs (bjerager et al. 2018a, this volume) for three selected intervals of the blokelv-1 core. the ankerite contents were determined by point counting of thin sections and show that ankerite-cemented beds are characterised by a bulk density of about 2.75 g/cm3. 75 petrographic relationships the detrital components in the sandstones comprise quartz, albite, k-feldspar, mica minerals, rock fragments, heavy minerals and organic matter (fig. 6). furthermore, macroscopic mudclasts and bioclasts are often present within the sandstones (fig. 2). elongate detrital grains such as mica minerals are generally oriented parallel to the layering, and the mica grains are often bent around more resistant grains. most of the organic matter in the pervasively ankerite-cemented sandstones consists of inertinite (fig. 10a). the authigenic components in the sandstones mainly comprise microquartz, quartz overgrowths, ankerite, kaolinite, illite, ti-oxides and pyrite. pyrite framboids are common in the sandstones and are typically present within organic matter. framboidal pyrite was also precipitated within muscovite grains prior to mica compaction. well-developed kaolinite booklets constitute the most common clay morphology (figs 10b, c). kaolinite is often present within and adjacent to albite grains and may fill large pore spaces (fig. 10d). kaolinite has grown locally between the cleavage planes of mica and beginning illitisation of kaolinite is observed (fig. 10e). kaolinite is often present around partly dissolved k-feldspar grains, but not within the secondary pores; such k-feldspar dissolution has only occurred in some of the k-feldspar grains (fig. 10f). illite is the only mineral that has precipitated within the secondary porosity in some of the k-feldspar grains (fig. 11a). microquartz often coats quartz and feldspar grains (fig. 11b) and is often enclosed in other authigenic phases such as ankerite and illite. quartz overgrowths are volumetrically small and the microquartz coatings apparently inhibited their formation. the quartz overgrowths are often interrupted and overgrown by illite (fig. 11c– d). illite was precipitated later than microquartz, both simultaneously with and later than ankerite. feldspar overgrowths are enclosed in the ankerite cement (fig. 11e). feldspars and muscovite have in some places been partly replaced by illite and mixed-layer clays. mixed-layer clays constitute various combinations of clay minerals and can be difficult to recognise. small amounts of ca and na are often found in the illite, however, and indicate the presence of mixed-layer clays such as smectite–illite. apatite has occasionally precipitated simultaneously with illite. small ti-oxide crystals have commonly formed simultaneously with illite within the clay (fig. 11c). precipitation of ti-oxides was contemporaneous with the formation of euhedral pyrite, but continued after pyrite formation ceased. euhedral pyrite crystals formed later than ankerite. initial stylolitisation is observed in places along the numerous mica grains and organic matter. ankerite cement mainly precipitated simultaneously with macroquartz (fig. 11c) and often formed around and within dissolving albite and plagioclase grains. the ankerite crystals are often zoned corresponding to an increasing fe-content from core to rim whereas the cacontent and especially the mg-content decrease. the ankerite cement is pervasive in some sandstone intervals, which are in general more fine-grained than the sandstones with less intensive cementation (table 1). the poikilotopic ankerite replaced some of the quartz, feldspar and mica grains partly or fully (fig. 11f). partial albitisation of some of the k-feldspar and plagioclase grains occurred prior to the poikilotopic ankerite phase. four igneous intrusions are present in the cored succession at depths of 100.10–102.04, 55.20–56.40, 26.40–27.10 and 7.05–7.35 m (fig. 2); they consist of sparsely plagioclase-olivine-phyric tholeiitic basalt (fig. 12a; larsen 2018, this volume). three were intruded in sandstone intervals where they have affected the diagenesis of the adjacent sandstones, albeit only at a centimetre-scale. the changes involve a second episode of microquartz precipitation that formed on surfaces of macroquartz and ankerite simultaneously with continued illite and ankerite precipitation (fig. 12b, c). resumed quartz overgrowth subsequently enveloped some of the microquartz. ankerite-filled fractures up to 0.5 mm thick are present near some of the sandstone–basalt contacts (fig. 12d). the fractures both cross-cut the contact or run parallel to the contact within the sandstone. the ankerite becomes more fe-rich towards the middle of the fractures where pyrite and cavities are sometimes present. patches of pyrite are locally abundant in the sandstones near the intrusions (fig. 12d). the groundmass of the basalts is partly altered in the chilled margins (fig. 12a). olivine is altered and partly dissolved within the chilled margins of the intrusions, whereas plagioclase is fresh (fig. 12e; larsen 2018, this volume). the ba-feldspar celsian was recorded immediately above the thickest intrusion where it forms abundant crystals with sizes of 5–10 µm (fig. 12f). kaolinite is rare to absent in sandstones next to intrusions, whereas it is the dominant clay mineral in all other sandstones (fig. 5). 7676 ka dq an mi qo ab mq ka ab kf il mq ik qo il ka qo il mi dq ti ka om an dq mi mol dia 10 10 µm20 µm 10 µm 10 µm 20 µm20 µm a b c d e f fig. 10. petrographic relationships in the hareelv formation sandstones. a: inertinite enclosed in ankerite cement. crossed nicols image, 151.71 m. b: pore-filling kaolinite booklets. secondary electron image, 69.78 m. c: growth of kaolinite, macroquartz, illite and ti-oxides. backscatter electron image, 7.41 m. d: albite partially replaced by kaolinite. backscatter electron image, 7.41 m. e: partial illitisation of kaolinite that is partly enclosed in macroquartz. secondary electron image, 221.18 m. f: partly dissolved k-feldspar grains enclosed in microquartz and illite. secondary electron image, 100.05 m (photo courtesy of morten l. hjuler). ab: albite. an: ankerite cement. dq: detrital quartz. ik: illitisation of kaolinite. il: illite. ka: kaolinite. kf: k-feldspar. mi: mica. mq: microquartz. om: organic matter. qo: quartz overgrowth. ti: ti-oxide. 77 discussion no systematic mineralogical differences are found between the in situ and remobilised, intruded gravity-flow deposits, so the post-depositional remobilisation of some of the sandstones (fig. 2) has not influenced the diagenesis. the similar mineralogical composition of the three deepest sandstone samples (fig. 5), which were sampled at the base, middle and top of a single gravity-flow bed, indicates that no significant mineralogical variation is present within such a sandstone body. these results further indicate that diagenesis in the sand layers was independent of the distance to the bed margins with adjacent mudstone layers, suggesting that there was limited transport of dissolved solids in the pore water. clay-mineral formation authigenic kaolinite is present in the sandstones, but its formation is enigmatic in these marine basinal deposits since interaction with flowing fresh or brackish water, undersaturated in salt, is necessary for its formation (bjørlykke 1998). some of the kaolinite may be detrital, having formed in deltaic environments before the sand was transported to the basin floor by gravity flows, but at least some of the kaolinite booklets are so well preserved that they are likely to have formed in situ (fig. 10c). it is not clear, however, how the authigenic kaolinite formed in this marine setting, especially as it is assumed to have formed shortly after deposition. kaolinite often formed in association with dissolution of albite grains, as testified by its location, filling large pore spaces next to partly dissolved albite grains (fig. 10d). there is a weak upwards-decreasing trend in clay content in the cored sandstones accompanied by an upwards-increasing albite content (figs 5, 6); this is also reflected by a pronounced upwards increase in the nacontent (bjerager et al. 2018a, this volume). it appears therefore that more kaolinite formed at the expense of albite in the lower part of the core relative to higher in the section. this may be explained by the gradual deepening of the depositional setting during the accumulation of the katedralen member (surlyk 2003; bjerager et al. 2018a, this volume), such that flushing of the gravity-flow sands with undersaturated water became progressively less likely as the water depth in the jameson land basin increased and the shoreline transgressed. it is not clear, however, how such meteoric water penetrated far into the subsurface beneath the marine basin although it is possible that the geometry of the sand bodies favoured such a process. the gravity-flow deposits of the katedralen member are interpreted to have been deposited in steep-sided gullies that terminated in sheet-sand splays on the basin floor (surlyk et al. 2007). the gully-fill sands encased in impermeable muds may have channelled downward flow of undersaturated surface water. an alternative explanation is that the kaolinite may have formed when much of the jameson land basin was subaerially exposed in the latest volgian (surlyk & noenygaard 2005). the accompanying flushing with meteoric water might have reached deep into the sediments if the porous sand beds and injectites formed conduits for the flow. it is unlikely that kaolinite formed during cenozoic uplift since the low permeabilities caused by sandstone cementation, at least locally, and mudstone compaction and cementation would have precluded sufficient flow of undersaturated water. the igneous intrusions emplaced in the early cenozoic would have constituted additional flow barriers. beginning illitisation of kaolinite is observed (fig. 10e) and illite sometimes formed in the partly dissolved k-feldspar grains (fig. 11a), indicating that illite and quartz precipitated at the expense of kaolinite and kfeldspar (bjørlykke 2015). the presence of mixed-layer clays containing small amounts of na and ca suggests that some of the illite may have transformed from smectite (fig. 5), and mixed-layer illite–smectite has been previously recorded from mudstones of the hareelv formation (lindgreen & surlyk 2000). ankerite formation the locally high ankerite content of up to 41 wt% (fig. 5) shows that a volumetrically significant carbonate source must have been present within the sandstones. the clear correlation between grain size and ankerite content shows that ankerite precipitation was favoured in sandstones with a very fine grain size (fig. 4, table 1), probably because more biogenic carbonate mud was deposited in the finer-grained sediments, which then sourced carbonate cementation during burial. mouldic ankerite-filled macropores are not observed in the sandstones so dissolution of larger carbonate clasts and fossils was probably not a significant source of carbonate. formation of carbonate cement at rather shallow depth prior to significant mechanical compaction is indicated by the locally high ankerite content, but the cement may first have precipitated as calcite and then have been replaced by ankerite at deeper burial. this seems plausible 7878 qo mq an mi dq df an qo dq il py ti an qo dq df fo qo il kf il dq qo mol dia 11 10 µm 20 µm 20 µm 20 µm 20 µm 20 µm a b c d e f fig. 11. petrographic relationships in the hareelv formation sandstones. a: illite precipitated within secondary porosity in k-feldspar. backscatter electron image, 100.05 m. b: quartz overgrowth on the central grain surrounded by microquartz crystals of various dimensions. secondary electron image, 43.16 m. c: growth of ankerite and macroquartz was restricted by illite precipitation. backscatter electron image, 7.41 m. d: macroquartz and illite have grown simultaneously. secondary electron image, 221.18 m. e: feldspar and quartz overgrowths enclosed in ankerite. crossed nicols image, 166.70 m. f. plagioclase, quartz and muscovite partly replaced by ankerite. crossed nicols image, 151.71 m. an: ankerite cement. df: detrital feldspar. dq: detrital quartz. fo: feldspar overgrowth. il: illite. kf: k-feldspar. mi: mica. mq: microquartz. py: pyrite. qo: quartz overgrowth. ti: ti-oxide. 79 fig. 12. diagenesis induced by enhanced heat flow and altered pore-fluid composition near basaltic sills and dykes in the hareelv formation sandstones. a: sandstone and basalt have been altered in the contact aureole. direct light image, 55.20 m. b: microquartz and illite have grown simultaneously. secondary electron image, 100.05 m. c: most ankerite grew before precipitation of microquartz and illite. backscatter electron image, 100.05 m. d: ankerite and pyrite precipitated in pores and fractures near the intrusion. backscatter electron image, 55.18 m. e: olivine was dissolved in the chilled edge of the basalt and ankerite was precipitated. backscatter electron image, 55.20 m. f: abundant celsian crystals formed locally near the intrusion. secondary electron image, 100.05 m. an: ankerite cement. ao: altered olivine. cc: celsian crystal. dq: detrital quartz. il: illite. mq: microquartz. pp: plagioclase phenocryst. py: pyrite. qo: quartz overgrowth. cc pp ao an qo mq dq il py dq mq il dq anpy basaltbasalt altered basalt altered basalt altered sandstone altered sandstone mol dia 12 a b c d e f 10 µm 20 µm 200 µm 30 µm 5 µm 500 µm 8080 since quartz and feldspar overgrowths are enclosed in the ankerite cement (fig. 11e) and because such a replacement was presumably only possible once iron and magnesium had been released by the smectite to illite transition (boles 1978). although growing quartz crystals may displace carbonate cement, the ankerite probably formed late during burial because the ankerite crystals, including those that formed on quartz overgrowths, often become more fe-rich towards their rim (fig. 11c). ca, mg and mn correlate very well in the pca model (fig. 7) suggesting that these elements reside primarily in ankerite and can pinpoint the cemented intervals. maximum burial depth the section removed by erosion is estimated to have been 2.8 km thick at the blokelv-1 borehole location based on apatite fission-track analyses (green & japsen 2018, this volume); such burial corresponds to maximum palaeotemperatures of c. 100–105°c in the studied succession with a presumed palaeogeothermal gradient of 30°c/km. this is in agreement with the general interpretation of the magnitude of cenozoic uplift of the jameson land basin area (mathiesen et al. 2000; hansen et al. 2001) and fits well with the diagenetic changes observed in the present study. the presence of kaolinite and mixed-layer clays sets limits on the maximum possible temperature that the sediments can have been subjected to, since kaolinite is unstable in the presence of k-feldspar at temperatures above c. 130°c (bjørlykke et al. 1986; bjørlykke 1998). the rather small amount of authigenic quartz in the studied sandstones (fig. 6) is a result of the porosity-preserving effect of the microquartz coatings and the limited residence time at deepest burial, which occurred prior to the onset of late eocene uplift (green & japsen 2018, this volume). bending of mica around harder detrital grains occurred during mechanical compaction (fig. 10a), but no grain crushing or pressure dissolution of quartz grains parallel to the layering have been observed. incipient stylolitisation along concentrations of mica grains and organic matter is present in some of the sandstones. stylolite formation in sandstones generally initiates at around 2.5 km of burial and well-developed stylolites are typically formed at depths of >3.5 km (bjørlykke et al. 1986). fully developed stylolites have been reported in jurassic sandstones of the vardekløft group in outcrops on traill ø in east greenland (baron & parnell 2007), but they may have formed in response to heating by igneous intrusions at a presumed burial depth of 2.0–2.5 km (therkelsen 2016). however, the incipient stylolitisation in the hareelv formation appears to be unrelated to igneous intrusions and has thus presumably developed in response to elevated temperatures during burial. intrusion-induced diagenesis two basalt sills and one dyke with thicknesses of up to 1.9 m were intruded into sandstone intervals in the cored succession during the early eocene (fig. 2; larsen 2018, this volume). one sill of 0.7 m thickness intruded into a mudstone interval in which the thermal maturity is significantly enhanced within a few metres of the intrusion (bojesen-koefoed et al. 2018, this volume). in the sandstones, intrusion-induced mineralogical changes are only evident up to a metre from the intrusions, and the enhanced heat flow caused by the intrusions may have been highly localised as the intrusions are so thin. ba-feldspar (celsian) can precipitate during hydrothermal activity and low-grade metamorphism (moro et al. 2001). thus, the restricted occurrence of celsian close to one of the igneous intrusions (fig. 12f) suggests that the sandstone was subjected to high temperatures immediately adjacent to the intrusion. microquartz usually forms under shallow burial and requires the pore fluids to be supersaturated relative to quartz ( jahren & ramm 2000). thus, the second episode of microquartz precipitation that occurred in the deeply buried sandstones near intrusions (figs 12b, c) resulted from silica supersaturation in the pore fluids induced by interaction with the intrusions. all olivine crystals have been altered in the chilled margins of the intrusions (fig. 12e; larsen 2018, this volume) and this alteration probably represents the source of the si-enrichment as ankerite is often observed within the dissolved olivine crystals. this fresh supply of iron and magnesium from dissolved olivine was presumably what promoted precipitation of ankerite and pyrite in the contact aureoles of the intrusions. precipitation of ankerite must have occurred rapidly in the fractures created by the intrusions before the fractures were closed by mechanical compaction (fig. 12d). the pore fluids became increasingly more fe-rich, as recorded by the outwards-increasing fe-content in zoned ankerite crystals. kaolinite is not present near the intrusions in contrast to all other sandstones in the core; this indicates that the locally high temperatures near the intrusions caused kao81 linite to be altered to illite, explaining the increased illite content near the intrusions despite a total clay content that is similar to sandstones elsewhere in the cored section. reservoir quality a tentative porosity development during burial is presented in fig. 13 concerning those diagenetic changes that most affected the porosity. the timing of the precipitation of the various minerals is determined by their petrographic relationships to known processes of mineral alteration. however, the timing of kaolinite precipitation and the presence and timing of an early calcite cement are uncertain. the porosity is high in the uncemented sandstones in the blokelv-1 core (fig. 8), but some of it is present as microporosity in low-permeable clay minerals. thus, the permeability ranges from intermediate to high, also because the sandstones are generally fine-grained (table 1; beard & weyl 1973). porosities >20% are only encountered in the upper 130 m of the core (bjerager et al. 2018a, this volume), most likely due to the upwardsdecreasing clay content in the sandstones (fig. 6). pervasive ankerite precipitation is the diagenetic process that has most significantly reduced porosity and permeability (figs 8, 9), but the largest amounts of ankerite formed in the sandstones that already at deposition had the lowest permeability due to their very fine grain size (table 1). the highest porosities are present in the most quartzrich sandstones (fig. 6) where microquartz coatings inhibited the precipitation of macroquartz. the sandstones were deeply buried for a limited period of time (fig. 13; green & japsen 2018, this volume), so the formation of quartz overgrowths and illite had not decreased reservoir quality critically (fig. 8). the igneous intrusions reduced the porosity significantly in the contact aureoles of the j c p e o m p 160 140 120 100 80 60 40 20 0time (ma) time period 20 4105(120)70 2.80 1 2 0 po ro sit y (% ) 0 20 10 30 40 150 130 110 90 70 50 30 10 eogenesis mesogenesis telogenesis mechanicalcompaction diagenetic regime temperature (°c) burial depth (km) microquartz macroquartz,illite fineand medium-grained sandstone very fine-grained sandstone sandstone near an igneous intrusion m icroquartz, ankerite, illite kaolinite intrusions c alcite macroquartz,ankerite, illite mechanicalcompaction microquartz mol dia 13 fig. 13. tentative porosity evolution and diagenetic development of the hareelv formation sandstones. the correlation between burial and time is derived from apatite fission-track analyses (green & japsen 2018, this volume); the timing of igneous intrusion is from larsen (2018, this volume). the mechanical compaction curve is based on a study of microquartz-coated sandstones (ramm et al. 1997); the onset of mesogenesis, where chemical compaction succeeds mechanical compaction, is estimated at about 70°c (morad et al. 2000). the presence and timing of early calcite cement is uncertain. the present-day porosity of the three sandstone classes illustrated is taken as an average of the analysed representatives. 8282 adjacent sandstones due to the mineralogical changes caused by the enhanced heat flux and altered pore-fluid composition (fig. 13). this is a strictly local phenomenon, however, since the influence is minimal in sandstones more than a metre away from the intrusions. most of the igv values give an unrealistically low estimate of the depositional porosity (fig. 6), which is mainly due to mechanical compaction (houseknecht 1987; ehrenberg 1989). in addition, porosity is underestimated during point counting when microporosity is present, as reflected by the higher values measured by core analysis (fig. 8; dutton & loucks 2010; olivarius et al. 2015). values of the calculated permeability log are overestimated in the mudstones (fig. 9) because the calculation of permeability is based on the total porosity including microporosity, which does not increase permeability. the cored sandstone layers are rather thin (fig. 2), but they presumably have good vertical connectivity since most of them are remobilised (bjerager et al. 2018a, this volume). thus, the reservoir quality of the sandstones is intermediate to good in the intervals that lack pervasive ankerite cement and abundant clay minerals. conclusions the dominant diagenetic changes that have occurred in the sandstones of the hareelv formation in the jameson land basin comprise precipitation of kaolinite, microquartz, ankerite, macroquartz and illite. the largest porosity reduction was caused by pervasive ankerite, which formed in sandstones that already at the time of deposition had the lowest permeability due to their very fine grain size. in general, the porosity of the remaining sandstones is high and the permeability is intermediate to high since microquartz coatings have inhibited macroquartz formation, although some of the sandstones have poorer reservoir quality due to the presence of clay minerals. igneous intrusions have lowered the porosity and permeability significantly in the closely adjacent sandstones where the primary diagenetic changes induced by the intrusions include renewed microquartz precipitation and accelerated formation of ankerite and illite. acknowledgements the authors wish to thank helene almind, kirsten fries, karen henriksen, anders pilgaard and john boserup for sampling and analysis. valuable advice by lotte m. larsen, tonci balic-zunic, holger lindgreen and morten bjerager is much appreciated as well as artwork by jette halskov and stefan sølberg. the referees, knut bjørlykke and jens therkelsen, are thanked for insightful comments that improved the manuscript significantly. references alsen, p. & piasecki, s. 2018: biostratigraphy of the hareelv formation (upper jurassic) in the blokelv-1 core, jameson land, central east greenland. in: ineson, j. & bojesen-koefoed, j.a. (eds): petroleum geology of the upper jurassic – lower cretaceous of east and north-east greenland: blokelv-1 borehole, jameson land basin. geological survey of denmark and greenland bulletin 42, 15–37 (this volume). baron, m. & parnell, j. 2007: relationships between stylolites and cementation in sandstone reservoirs: examples from the north sea, u.k. and east greenland. sedimentary geology 194, 17–35. beard, d.c. & weyl, p.k. 1973: influence of texture on porosity and permeability of unconsolidated sand. aapg bulletin 57, 349–369. bjerager, m., alsen, p. bojesen-koefoed, j.a., piasecki, s. & pilgaard, a. 2018b: late jurassic evolution of the jameson land basin, east greenland – implications of the blokelv-1 borehole. in: ineson, j. & bojesen-koefoed, j.a. 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(eds): petroleum geology of the upper jurassic – lower cretaceous of east and north-east greenland: blokelv-1 borehole, jameson land basin. geological survey of denmark and greenland bulletin 42, 39–64 (this volume). bjørlykke, k. 1998: clay mineral diagenesis in sedimentary basins – a key to the prediction of rock properties. examples from the north sea basin. clay minerals 33, 15–34. bjørlykke, k. 2015: petroleum geoscience. from sedimentary environments to rock physics, 650 pp. berlin, heidelberg: springer verlag. bjørlykke, k., aagaard, p., dypvik, h., hastings, d.s. & harper, a.s. 1986: diagenesis and reservoir properties of jurassic sandstones from the hahenbanken area, offshore mid norway. in: spencer, a.m. et al. 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(eds): petroleum geology of the upper jurassic – lower cretaceous of east and north-east greenland: blokelv-1 borehole, jameson land basin. geological survey of denmark and greenland bulletin 42, 85–113 (this volume). 83 boles, j.r. 1978: active ankerite cementation in the subsurface eocene of southwest texas. contributions to mineralogy and petrology 68, 13–22. dutton, s.p. & loucks, r.g. 2010: reprint of: diagenetic controls on evolution of porosity and permeability in lower tertiary wilcox sandstones from shallow to ultradeep (200–6700 m) burial, gulf of mexico basin, u.s.a. marine and petroleum geology 27, 1775–1787. ehrenberg, s.n. 1989: assessing the relative importance of compaction processes and cementation to reduction of porosity in sandstones: discussion; compaction and porosity evolution of pliocene sandstones, ventura basin, california: discussion. aapg bulletin 73, 1274–1276. esbensen, k. 2010: multivariate data analysis – in practice. 5th edition, 598 pp. oslo: camo software. green, p.f. & japsen, p. 2018: burial and exhumation history of the jameson land basin, east greenland, estimated from thermochronological data from the blokelv-1 core. in: ineson, j. & bojesenkoefoed, j.a. 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(eds): petroleum geology of the upper jurassic – lower cretaceous of east and north-east greenland: blokelv-1 borehole, jameson land basin. geological survey of denmark and greenland bulletin 42, 115–126 (this volume). ramm, m., forsberg, a.w. & jahren, j.s. 1997: porosity-depth trends in deeply buried upper jurassic reservoirs in the norwegian central graben: an example of porosity preservation beneath the normal economic basement by grain-coating microquartz. in: kupecz, j.a., gluyas, j. & bloch, s. (eds): reservoir quality prediction in sandstones and carbonates. american association of petroleum geologists memoir 69, 177–199. rietveld, h.m. 1969: a profile refinement method for nuclear and magnetic structures. journal of applied crystallography 2, 65–71. slater, c. & cohen, l. 1962: a centrifugal particle size analyser. journal of scientific instruments 39, 614–617. stemmerik, l., christensen, f.g., piasecki, s., jordt, b., marcussen, c. & nøhr-hansen, h. 1992: depositional history and petroleum geo logy of the carboniferous to cretaceous sediments in the northern part of east greenland. norwegian petroleum federation, special publication 2, 67–87. surlyk, f. 2003: the jurassic of east greenland: a sedimentary record of thermal subsidence, onset and culmination of rifting. in: ineson, j.r., surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 659–722. surlyk, f. & noe-nygaard, n. 2005: a forced regressive shelf-margin wedge formed by transition-slope progradation: lowermost cretaceous rauk plateau member, jameson land, east greenland. bulletin of the geological society of denmark 52, 227–243. surlyk, f., gjelberg, j. & noe-nygaard, n. 2007: the upper jurassic 8484 hareelv formation of east greenland: a giant sedimentary injection complex. in: hurst, a. & cartwright, j. (eds): sand injectites: implications for hydrocarbon exploration and production. aapg memoir 87, 141–149. therkelsen, j. 2016: diagenesis and reservoir properties of middle jurassic sandstones, traill ø, east greenland: the influence of magmatism and faulting. marine and petroleum geology 78, 196–221. _________________________________________________________________________________________ manuscript received 10 june 2016; revision accepted 29 august 2017 geological survey of denmark and greenland bulletin 22, 2010, pp. 92 geological survey of denmark and greenland bulletin 22· 2010 lithostratigraphy of the upper oligocene – miocene succession of denmark erik skovbjerg rasmussen, karen dybkjær and stefan piasecki geological survey of denmark and greenland ministry of climate and energy bulletin 22_ gsb191-indhold 04/03/11 12.40 side 1 geological survey of denmark and greenland bulletin 22 keywords lithostratigraphy, miocene, upper oligocene, north sea, denmark cover upper shoreface sands of the billund formation in the addit gravel pit, central jylland showing prominent burrows (ophiomorpha isp.); illustrated section is 40 cm high. photo: tom pallesen. frontispiece outcrop of the kolding fjord member, klintinghoved formation at hagenør, lillebælt. the succession is dominated by organic-rich, lagoonal, silty clay and sand beds deposited as washover fans on the back-barrier flat. photo: peter warna-moors. chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, department of earth sciences, university of aarhus; minik rosing, geological museum, university of copenhagen; finn surlyk, department of geography and geology, university of copenhagen scientific editor of this volume: jon r. ineson editorial secretaries: jane holst and esben w. glendal referees: dan evans (uk) and claus heilmann-clausen (dk) illustrations: stefan sølberg digital photographic work: benny m. schark layout and graphic production: henrik klinge pedersen printers: rosendahls · schultz grafisk a/s, albertslund, denmark manuscript received: 15 december 2009 final version approved: 1 november 2010 printed: 31 december 2010 issn 1604-8156 isbn 978-87-7871-291-2 citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 22, 92 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk or at www.geus.dk/publications/bull © de nationale geologiske undersøgelser for danmark og grønland (geus), 2010 for the full text of the geus copyright clause, please refer to www.geus.dk/publications/bull bulletin 22_ gsb191-indhold 04/03/11 12.40 side 2 bulletin 22_ gsb191-indhold 04/03/11 12.40 side 3 4 contents abstract . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5 introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7 geological setting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9 previous studies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 11 previous lithostratigraphic subdivision . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 11 vejle fjord formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 11 klintinghoved, ribe and arnum formations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 12 arnum, hodde, gram and sæd formations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 12 odderup formation (terrestrial miocene) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 12 sequence stratigraphy and onshore–offshore correlation . . . . . . . . . . . . . . . . . . . . . . . . . . 13 data and methodology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 14 revised lithostratigraphy . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16 brejning formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16 sydklint member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 20 øksenrade member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 21 ribe group . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 24 vejle fjord formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 27 skansebakke member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 32 billund formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 34 hvidbjerg member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 36 addit member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 40 klintinghoved formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 45 kolding fjord member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 50 bastrup formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 53 resen member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 55 arnum formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 56 vandel member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 58 odderup formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 59 stauning member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 61 fasterholt member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 63 måde group . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 65 hodde formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 66 ørnhøj formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 68 gram formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 70 marbæk formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 72 stratigraphic architecture . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 75 palaeogeography . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 80 acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 87 references . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 88 bulletin 22_ gsb191-indhold 04/03/11 12.41 side 4 5 abstract authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: esr@geus.dk *present address: university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. rasmussen, e.s., dybkjær, k. & piasecki, s*. 2010: lithostratigraphy of the upper oligocene – miocene succession of denmark. geological survey of denmark and greenland bulletin 22, 92 pp. this paper presents a revised lithostratigraphic scheme for the uppermost upper oligocene – miocene succession of denmark. the marine oligocene brejning clay member is upgraded to formation status and includes the sydklint member and the øksenrade member (new). the shallow marine and deltaic deposits of mainly early miocene age are included in the ribe group (new) while the fully marine middle and upper miocene clay-rich deposits are referred to the måde group (new). the ribe group is subdivided into 6 formations: the vejle fjord formation is revised and includes the skansebakke member, the billund formation (new) includes the addit and hvidbjerg members (new), the klintinghoved formation is redefined formally and includes the koldingfjord member (new), the bastrup formation (new) includes the resen member (new), the vandel member is a new member in the arnum formation (revised), the odderup formation is redefined and includes the stauning member (new) and the coalbearing fasterholt member. the måde group is subdivided into the hodde, ørnhøj (new), gram and marbæk (new) formations. subdivision of the upper oligocene – miocene succession into two groups, the ribe and måde groups, is compatible with the north sea lithostratigraphic framework where they correlate with the upper part of the hordaland group and the nordland group, respectively. the revised lithostratigraphic framework correlated in three dimensions provides rigorous constraints on the palaeogeographic interpretation of the late oligocene – miocene period. three major deltaic units (billund, bastrup and odderup formations) prograded from the north and north-east into the north sea basin during the early – early middle miocene. delta progradation was punctuated by deposition of marine clay and silt associated with minor transgressive events (vejle fjord, klintinghoved and arnum formations). during the middle–late miocene, marine depositional conditions dominated (hodde, ørnhøj and gram formations). a fourth and final progadational event (marbæk formation) commenced in the latest tortonian heralding the emergence of present-day denmark (including the north sea sector). bulletin 22_ gsb191-indhold 04/03/11 12.41 side 5 6 fig. 1. map showing the location of the study area in jylland, western denmark, and the boreholes, outcrops and seismic data used. towns and villages mentioned in the text are indicated. the index map (lower right) shows the localities in the fredericia–kolding area. 56°n fig. 77 fig. 78 57°n 56°n 55°n 8°e 10°e fig. 79 borehole outcrop town seismic lines seismic sections shown in figures 25 km voervadsbro sdr. vissing århussilkeborg føvling rømø sylt ribe ribe gram arnum-1 rødding estrup bastrup lillebælt hvidbjergpjedsted vejle fjord horsens fjord limfjorden kvong forumlund marbæk sjelborg måde esbjerg borg-1 løgumkloster vollerup vester sottrup hørup havsønderborg klintinghoved tinglevsæd v. torsted lundgård hodde vorbasse grindsted billund billund vandel gadbjerg almstok egtved andkær skansebakke vejle horsens søvind juelsmindesanatoriet fakkegrav brejning brejning hoved jensgård dykær stakrogeodderupskjern assing mølleby hammerum herning isenvad fasterholt søby store vorslunde give brande hjøllund fjelstervang fjand ulfborg ørnhøj, lille spåbæk abildå holstebro klosterhede vind stensig stauning sdr. vium resen brøndum lyby mogenstrup gyldendal søndbjerg lodbjerg skyum skyum bjerge skanderup thisted silstrup klovbakker mors sunds ikast bording mausing sorring sofienlund hinge ølst gl. rye morsholt salten addit addit mark uldum lindved hovslund hellevad struer skive viborg vonsild kolding fjord flensborg fjord lillebæ lt middelfart fredericia øksenrade fænø hindsgavl galsklint kolding røjle klint børup rønshoved hagenør lillebælt north sea uk norway sweden germany poland denmark bulletin 22_ gsb191-indhold 04/03/11 12.41 side 6 7 sediments of miocene age crop out in denmark along the east coast of jylland and in the limfjorden area (fig. 1); a few inland cliffs in central jylland also expose miocene deposits, especially in river scars and road cuts. excavation for raw materials for construction, i.e. gravel, sand and clay, has resulted in open pits that expose miocene deposits, mainly in western and central jylland. during the last decade, the increasing need for water for domestic purposes and irrigation has initiated intensive drilling programs and the acquisition of high-resolution seismic data from the miocene succession. the renewed interest in the miocene has resulted in financial support for field investigations, so it has been possible to re-study all danish outcrops exposing miocene deposits. a high-resolution biostratigraphic subdivision of the miocene succession has been developed, based on dinoflagellate cysts (dinocysts) (dybkjær & piasecki 2010). this new dinocyst zonation has provided a robust framework within which the studied boreholes have been correlated. it has also made it possible to integrate all seismic and borehole data with the new, detailed sedimentological descriptions and interpretations of the outcrops (friis et al. 1998; rasmussen & dybkjær 2005; e.s. rasmussen et al. 2006) in order to construct a depositional model for the miocene succession. associated studies, for example of the climatic conditions (larsson et al. 2006; larsson-lindgren 2009; t. utescher, personal communication 2009) and sediment provenance (knudsen et al. 2005; olivarius 2009), have further added to the understanding of the depositional system. the lithostratigraphy presented here encompasses the upper upper oligocene – miocene succession found onshore denmark. it is bounded beneath by a major unconformity between upper eocene – lower upper oligocene clay-rich deposits and siltand sand-rich deposits of late late oligocene – miocene age. the top of the successsion is defined by the quaternary unconformity. during the study of the succession, it was necessary to establish a number of informal lithostratigraphic units that are now widely used in the mapping of aquifers both in denmark and germany, and are increasingly adopted in the literature (rasser et al. 2008; knox et al. 2010). it is therefore timely to formally define these units and redefine existing lithostratigraphic units in order to construct a consistent lithostratigraphic framework. the miocene succession was deposited during a period of worldwide tectonism (potter & szatmari 2009) and marked climatic change (e.g. zachos et al. 2001; miller et al. 2005; utescher et al. 2009). two of the most distinct phases in the alpine orogeny commenced in the miocene, the late oligocene – early miocene savian phase and the middle miocene betic phase (ziegler 1982; oszczypko 2006; ribero et al.1990). the opening of the north atlantic was characterised by the final change in spreading from the aegir ridge to the kolbeinsey ridge and increasing spreading rates in the early miocene have been detected (mosar et al. 2002; doré et al. 2008). in the middle miocene, a major tectonic reorganisation occurred (ziegler 1982; doré et al. 2008). the climate was warm temperate in the early – early middle miocene, but changed to a cold temperate climate in the late miocene. the miocene succession studied here was deposited in the eastern part of the north sea basin (fig. 2). the onshore portion of this basin under focus here is a stratigraphically complete fluvial – deep shelf transect that is recorded in detail by outcrop, borehole and seismic data; it provides a natural laboratory for the study of the development of fluviodeltaic depositional systems, the tectonic impact on basin evolution and the consequences of climatic changes including glacio-eustatic sea-level changes. in addition to creating a robust and consistent framework for practical applications, therefore, the lithostratigraphic revision presented here is a prerequisite for future research into miocene climatic, tectonic and eustatic evolution. introduction bulletin 22_ gsb191-indhold 04/03/11 12.41 side 7 8 fig. 2. palaeogeographic reconstruction of north-west europe during the early miocene (modified from e.s. rasmussen et al. 2008); configuration based on mosar et al. (2002). 100 km bulletin 22_ gsb191-indhold 04/03/11 12.41 side 8 9 the evolution of the north sea basin was strongly influenced by the collision between the african and european tectonic plates, volcanism in central europe and the opening of the north atlantic (ziegler 1982; 1990; ziegler et al. 1995; martinsen et al. 1999; faleide et al. 2002; e.s. rasmussen et al. 2005, 2008; rasmussen 2009a; gabrielsen et al. 2010). interaction of these factors with changing eustatic sea level resulted in final closure of the southern connection with the tethyan ocean during early–middle miocene times (harzhauser & piller 2007); subsequently, the only connection to the atlantic was through a strait between norway and shetland (fig. 2). the depositional basin of the eastern north sea area which covered present-day denmark, was bounded towards the north-east by the fennoscandian shield (fig. 3; bertelsen 1978; vejbæk 1997). the transition to the basin was controlled by the se–nw-trending sorgenfrei–tornquist zone. the basin was subdivided into two subbasins: the norwegian –danish basin and the north german basin, with the ese–wnw-striking ringkøbing–fyn high separating the subbasins. the ringkøbing–fyn high is further segmented into a number of n–s-trending elements such as the brande trough (fig. 3). these structural elements were formed during permian rift tectonics and later reactivated in the jurassic and during late cretaceous and early paleocene inversion tectonics (ziegler 1990; liboriussen et al. 1987; mogensen & jensen 1994; vejbæk & andersen 2002). reactivation of some of the older structures occurred in the oligocene as well as in the miocene (e.s. rasmussen 2004a, 2009a; japsen et al. 2007). during the middle miocene, the north sea basin experienced increased regional subsidence (ziegler 1982, 1990; vejbæk 1992; koch 1989; michelsen et al. 1998; clausen et al. 1999; e.s. rasmussen 2005). in the late pliocene – early plei stocene, the north sea basin was tilted towards the southwest (japsen 1993; japsen & bidstrup 1999; japsen et al. 2002; e.s. rasmussen et al. 2005). the north sea basin was located in the northern westerly wind belt. the climate was warm temperate to tropical in the early part of the paleogene (buchardt 1978; heilmann-clausen & surlyk 2006; zachos et al. 2001). a dramatic change occurred at the eocene–oligocene transition where a distinct climatic cooling took place. the early oligocene icehouse climate resulted in a marked eustatic sea-level drop due to growth of ice caps, primarily on antartica (buchardt 1978; prentice & matthew 1988; miller et al. 1991, 1996, 1998, 2005; zachos et al. 2001). however, by the end of the oligocene a subtropical climate prevailed in the north sea basin area (t. utescher, personal communication 2009; larsson et al. 2010). at the boundary between the palaeogene and the neogene, a marked, but transient, climatic cooling occurred with buildup of widespread ice caps on antarctica. this climatic event resulted in a major, global sea-level fall (miller et al. 1998; zachos et al. 2001). the early miocene climate in the north sea basin area was characterised by an oscillation between cool temperate and warm temperate climates (mai 1967; larsson et al. 2006). an overall increase in temperature culminated at the early to middle miocene transition, the so-called ‘mid-miocene climatic optimum’ (buchardt 1978; zachos et al. 2001). in the north sea basin area, a warm temperate to subtropical climate prevailed (mai 1967; friis 1975; utescher et al. 2000, 2009). geological setting fig. 3. structural elements in the study area. modified from berthelsen (1992). 57°n 56°n 55°n 58°n 8°e 12°e10°e norway denmark sweden germany ringkøbing – fyn high sorgenfrei–tornquist zone 50 km faults positive structural elements norwegian–danish basin fennoscandian shield rødd ing g ra be n brande trough tønder graben bulletin 22_ gsb191-indhold 04/03/11 12.41 side 9 10 at the termination of the middle miocene, a marked drop in global temperature commenced and during most of the late miocene the north sea basin area was characterised by a cool temperate climate (buchardt 1978; utescher et al. 2000, 2009; zachos et al. 2001; larsson-lindgren 2009). fine-grained siciliclastic sediments of mainly deep marine origin were deposited in denmark and the north sea basin during the post-danian palaeogene (heilmann-clausen et al. 1985; heilmann-clausen 1995; schiøler et al. 2007). a general sea-level lowstand and tectonic re-organisation during the oligocene resulted in erosion or non-deposition, especially in the central and southern part of the study area. in the northern part of the north sea basin, prodeltaic, clay-dominated wedges were laid down. in the latest oligocene, renewed transgression resulted in the deposition of glaucony-rich clay. this was followed by deposition of deltaic and coastal-plain sand and clay in the early miocene. three major deltaic progradational pulses occurred during the early miocene; the third and final pulse was characterised by extensive coal deposition. subsequent to deposition of the dominantly fluvio-deltaic deposits in the early to early middle miocene, full marine, clay-rich sedimentation dominated during the remaining part of the middle and late miocene. late pliocene – early pleistocene tilting of the eastern north sea area (japsen 1993; japsen & bidstrup 1999; japsen et al. 2002; e.s. rasmussen et al. 2005), combined with periodic growth of ice caps in the northern hemisphere, resulted in base-level fall and marked erosion of the substratum; middle and upper miocene deposits are thus missing in the eastern and northern parts of jylland (fig. 4). lower miocene oligocene eocene middle–upper paleocene lower paleocene (danian) upper cretaceous lower cretaceous and older middle–upper miocene 50 km57°n 56°n 55°n 8°e 12°e 10°e fig. 4. pre-quaternary subcrop map of denmark. modified from sorgenfrei & berthelsen (1954) and håkansson & pedersen (1992). bulletin 22_ gsb191-indhold 04/03/11 12.41 side 10 11 j.g. forchhammer (1794–1865) wrote the first account of the geology of denmark (forchhammer 1835; see also garboe 1961) and described the diluvial ‘rullestens dannelse’ (loosely translated as ‘boulder formation’) which he recognised throughout denmark; this unit was partly included in the tertiary. the lower part was named the ‘amber–brown-coal formation’ (translated from danish) and included fossiliferous strata of ‘the western system’, which was recognised in west and central jylland as well as farther south in germany, for example on the island of sylt (fig. 1). this ‘western system’ undoubtedly included the marine miocene as recognised today. beyrich (1853) studied molluscs collected by forchhammer from sylt and informed forchhammer in 1854 that he had identified these as a miocene fauna (garboe 1961). molluscs from south-west jylland (e.g. esbjerg and gram) were also identified as being of miocene age, and the results were presented at the 11th scandinavian research meeting in copenhagen in 1873 (mørch 1874). the palaeontologist j.p.j. ravn (1866–1951) established the first miocene (and oligocene) stratigraphy of denmark based on his comprehensive study of the fossil faunas in dark brown and grey, mica-rich clay which occurred widely in jylland. the resulting stratigraphic scheme of the lower, middle and upper miocene and associated deposits (ravn 1906) was published one year before his monograph on the oligocene and miocene mollusc faunas (ravn 1907). ravn realised that lower miocene marine faunas were missing and therefore suggested that the widespread brown-coal deposits represented the lower miocene. he also included part of the mica-rich clay and sand succession of the lillebælt region in south-east jylland in the lower miocene, based on mixed oligocene–miocene faunas. the botanist n.e.k. hartz (1867–1937) studied the succession that includes brown-coal deposits. exposures of brown coals were scarce at that time, but he concluded that the coals and the associated mica-rich sediments are all freshwater deposits (hartz 1909) and he found no evidence to contradict the early miocene age suggested by ravn (1906). the geological survey of denmark (dgu) performed two drilling campaigns in 1917 and 1921 under the leadership of v. milthers, and more brown-coal deposits were located. later, on the initiative of k. milthers, dgu drilled almost 9000 boreholes during the years between 1941 and 1949 (l.b. rasmussen 1988). in the last campaign (1958–1963), more than 2000 boreholes were drilled, making a total of approximately 11 000 boreholes (l.b. rasmussen 1988). in addition to engineering data, these extensive programmes also yielded geological results such as the volume, numbers and extent of brown-coal seams. the middle–upper miocene succession was not well understood prior to the second world war but the second drilling campaign revealed the overall stratigraphy and approximate thickness of these strata (milthers 1949; heller 1960). these workers proposed the existence of two discrete coalbearing units (probably broadly equivalent to the ribe and odderup formations of later workers, see below) underlying c. 100 m of middle miocene marine sediments (the arnum formation of later workers). previous lithostratigraphic subdivision prior to this study, the upper oligocene – miocene succession of onshore denmark had been subdivided lithostratigraphically into a number of formal and informal units. the origin of these terms is discussed briefly below to provide the background to the lithostratigraphic revision presented in this bulletin. vejle fjord formation the mixed oligocene–miocene fauna reported by ravn (1907) from the lillebælt region (fig. 1) impelled eriksen (1937) to study the same succession in this region for fossils; he found a sparse mollusc fauna in the brejning exposure on the south coast of vejle fjord and in neighbouring cliffs. the fauna in the lower, glauconitic strata was of oligocene age, but the sparse fauna in the overlying black, micaceous clay was suggested to be of early miocene age. the uppermost beds of the succession, comprising micaceous, grey sand, were barren of fossils. larsen & dinesen (1959) studied the same strata in two exposures and formally established the vejle fjord formation which consisted of the glauconitic brejning clay and the overlying, black to grey, micaceous clay and sand of the vejle fjord clay and vejle fjord sand, respectively. analysis of the foraminifer fauna in the brejning clay clearly indicated an upper oligocene affinity whereas a contrasting foraminifer fauna in the overlying, basal vejle fjord clay was suggested to indicate an early miocene age (larsen & dinesen 1959). these data supported the results based on the mollusc fauna, indicating that most of the formation, i.e. the vejle previous studies bulletin 22_ gsb191-indhold 04/03/11 12.41 side 11 12 fjord clay and vejle fjord sand”, should be referred to the lower miocene. the oligocene–miocene transition was thus placed near the shift from the glauconitic brejning clay to the black pyritic clay of the vejle fjord clay. in the following years, the vejle fjord clay and vejle fjord sand were systematically excluded from most mio cene stratigraphic schemes (e.g. l.b. rasmussen 1961). it was not until much later that danish stratigraphers incorporated the vejle fjord formation and the foraminifer stratigraphy in a miocene stratigraphic scheme, although maintaining the vejle fjord formation and klintinghoved formation (see below) as separate geographic entities (buchardt-larsen & heilmann-clausen 1988). farther north in jylland, christensen & ulleberg (1973) defined the sofienlund formation which was divided into four members: the ulstrup clay, the sofienlund clay, the sofienlund silt and the sofienlund sand. the foraminifer content of the sofienlund formation suggested a chattian age for the two lower members and a post-chattian age for the upper two members (christensen & ulleberg 1973). the lithology and biostratigraphy clearly indicate that these sediments should have been referred to the previously established vejle fjord formation. a similar view was stated by larsen & kronborg (1994), according to whom the lower two members are equivalent to the brejning clay whereas the upper two members equate to the vejle fjord clay and vejle fjord sand. the sofienlund formation is abandoned herein. the sydklint member was erected in north-west jylland and provisionally referred to the basal vejle fjord formation by heilmann-clausen (1997); this member is re-assigned to the brejning formation in this bulletin. klintinghoved, ribe and arnum formations the fossil mollusc fauna of the lower and middle miocene, exemplified by fauna from the coastal cliff at klintinghoved and seven deep wells in southern jylland, was studied by sorgenfrei (1940, 1958). the klintinghoved formation (of present usage) was not defined formally, but arose from extensive palaeontological work on the outcrop of a glacial, dislocated and folded raft of sediments that was considered to be of early miocene age (sorgenfrei 1940). sorgenfrei (1957) included the ‘klintinghoved mica clay’ as a formal formation in his ‘formations of denmark’ in lexique stratigraphique. two new formations were defined on the basis of the deep wells, the ribe and arnum formations; the marine clay of the arnum formation was referred to the middle miocene on the basis of the fauna (sorgenfrei 1958). the faunally barren ribe formation, composed of quartzitic sand, was recorded below the fossiliferous arnum formation in one well near the town of ribe. in the danish american prospecting company (dapco) well at arnum, sorgenfrei (1958) tentatively referred quartz-rich gravel and sand, below mud and sand of the arnum formation, to the ribe formation and underlying clays to the klintinghoved formation based solely on the lithological succession, in the absence of a mollusc fauna. arnum, hodde, gram and sæd formations in the comprehensive stratigraphic work by l.b. rasmussen (1958, 1961, 1966, 1968), focus was on the upper arnum formation, and the hodde and gram formations; the latter two formations were formally erected (l.b. rasmussen 1961). he continued and extended sorgenfrei’s work, producing a biostratigraphic zonation of this stratigraphic interval. l.b. rasmussen (1966) referred the gram formation (including the silt interval encountered in the borehole at sæd) to the upper miocene and the hodde formation to the middle miocene (fig. 5), assignments that have been largely confirmed by later work. l.b. rasmussen (1961; see also laursen et al. 1998) suggested that sandy strata overlying the gram clay in south-west jylland could be of messinian (latest miocene) age based on a mollusc fauna that was considered to be incompatible with the upper gram clay faunas; the sæd formation, overlying the gram formation, was thus proposed. hinsch (1990) re-evaluated this mollusc fauna, however, demonstrating equivalence to the tortonian fauna in the uppermost gram formation; this is supported by dinoflagellate floras in the same strata (piasecki 2005). the sæd formation is therefore abandoned in the present paper. odderup formation (terrestrial miocene) the odderup formation was erected by l.b. rasmussen (1961) as the brown-coal and quartz-sand succession between the marine clays of the arnum formation and the overlying hodde formation. the observation of brown-coal or coal fragments and quartz sands below the marine arnum formation in certain wells, however, complicated the stratigraphic concept, but l.b. rasmussen (1961) envisaged two major prograding deltaic units (ribe and odderup formations) subdividing the marine miocene into three major units (klintinghoved formation, arnum formation, hodde–gram formations, see fig. 5). the geology of the søby–fasterholt area was published by koch (1989) in a comprehensive resumé of palaeo bulletin 22_ gsb191-indhold 04/03/11 12.41 side 12 13 botanical, sedimentological and stratigraphic studies, including an analysis of the brown-coal seams. in this area, the odderup formation is bounded by marine strata of the arnum formation beneath and the overlying marine succession of the hodde and gram formations. that part of the odderup formation containing brown coals was defined as the fasterholt member. sequence stratigraphy and onshore–offshore correlation in a study of the cenozoic of the danish north sea, michelsen (1994; michelsen et al. 1998) divided the late palaeogene–neogene succession into 3 allostratigraphic units: units 5 to 7. the succession was further subdivided into 11 depositional sequences. the unconformities recognised in the offshore geophysical data were not directly correlated onshore using seismic data, but were correlated to the onshore lithostratigraphic units based on the bio stratigraphic data available at that time. the danish offshore stratigraphy was integrated with the uk and nor we gian stratigraphy: units 5 and 6 were correlated with the upper hordaland group (lark formation of schiøler et al. 2007) whilst unit 7 was correlated with the nordland group. sequence stratigraphy was applied to the onshore mio cene succession in southernmost jylland based on the analysis of petrophysical logs from 6 wells combined with seismic data (e.s. rasmussen 1996). the succession was divided into 6 depositional sequences ranging in age from the latest oligocene to the latest miocene. precise dating of these sequences was precluded by a general lack of biostratigraphic data but the sequence stratigraphic framework was correlated with the existing lithostratigraphy. e.s. rasmussen (2004b) introduced a new sequence stratigraphic subdivision, this time based on 16 new boreholes, outcrops and multichannel seismic data distributed in central and southern jylland. this resulted in subdivision of the upper oligocene – miocene succession into 6 depositional sequences, a framework similar to that of e.s. rasmussen (1996), although the ages of the sequences were refined on the basis of dinoflagellate cyst stratigraphy that was established over this period and formalised recently (piasecki 1980, 2005; dybkjær & rasmussen 2000, 2007; dybkjær 2004a, b; dybkjær & piasecki 2008, 2010). fig. 5. miocene lithostratigraphy of western denmark as defined by l.b. rasmussen (1961). hodde formation arnum formation klintinghoved formation upper miocene middle miocene lower miocene lithostratigraphy gram formation west east chronostratigraphy marine deposits fluvio-deltaic deposits odderup formation ribe formation bulletin 22_ gsb191-indhold 04/03/11 12.41 side 13 14 data and methodology 10 15 20 a b nn12 n n 1 1 nn10 nn9 nn8 nn6 nn5 nn4 nn3 nn2 nn1 np25 nn7 l at e m id d le e ar ly m io ce n e oligocene a ge ( m a) epoch stage (ages in ma) n an n o p la n kt o n zo n at io n dinoflagellate cysts zonation: denmark (dybkjær & piasecki 2010) dinoflagellate events zonation messinian tortonian serravallian langhian burdigalian aquitanian chattian 23.03 20.43 15.97 13.65 11.61 7.25 5.33 amiculosphaera umbracula barssidinium evangelinae selenopemphix armageddonensis hystrichosphaeropsis obscura palaeocystodinium spp. gramocysta verricula achomosphaera andalousiense unipontidinium aquaeductum unipontidinium aquaeductum systematophara spp. cannosphaeropsis passio distatodinium biffii deflandrea phosphoritica, common chiropteridium galea caligodinium amiculum thalassiphora pelagica thalassiphora rota cordosphaeridium cantharellus exochosphaeridium insigne homotryblium spp. abundant labyrinthodinium truncatum labyrinthodinium truncatum cousteaudinium aubryae cousteaudinium aubryae ectosphaeropsis burdigalensis exochosphaeridium insigne sumatradinium hamulatum palaeocystodinium miocaenicum palaeocystodinium miocaenicum cerebrocysta poulsenii h. obscura (h. o.) g. verricula (g. v.) a. andalousiense (a. a.) u. aquaeductum (u. a.) l. truncatum (l. t.) a. umbracula (a. u.) c. cantharellus (c. c.) e. insigne (e. i.) c. aubryae (c. au.) c. galea (c. g.) d. phosphoritica (d. p.) s. hamulatum (s. h.) t. pelagica (t. p.) c. amiculum (c. am.) homotryblium spp. (h. spp) maximum occurrence first stratigraphic occurrence last stratigraphic occurrence fig. 6. dinocyst zonation for the uppermost oligocene – miocene succession onshore denmark, from dybkjær & piasecki (2010). the ages of the stage boundaries are from gradstein et al. (2004), the nannoplankton zonation from martini (1971). nn: neogene nannoplankton zone. np: palaeogene nannoplankton zone. dinoflagellate events indicated in black define zone boundaries, those indicated in grey are additional diagnostic events. twenty-five outcrops, one cored borehole at sdr. vium (dgu no. 102.948) and c. 50 boreholes, drilled using the airlift drilling technique, were available for the study (fig. 1). most of the boreholes were drilled in order to solve stratigraphic problems, but some were drilled in order to test seismic facies interpretations. all boreholes are identified by their dgu borehole numbers, whereas outcrops are referred to by the nearest locality name. all 25 outcrops and the cored borehole were described sedimentologically and samples taken for biostratigraphy. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 14 15 the grain size and mineralogy of the airlift borehole samples, each representing one metre, were described. in addition, c. 40 samples per borehole were taken for bio stratigraphic (dinocyst) analysis. the description of the fasterholt member, including the sedimentary logs, is based on koch (1989). in boreholes drilled using the airlift drilling technique, problems are experienced in retrieving fine-grained sand to the surface, and thus the recovery is commonly low or even zero in such intervals. as an aid to lithological identification, however, a gamma-ray log was obtained from all 50 boreholes. this petrophysical log is typically used to differentiate between sand and clay in siliciclastic sections, although sands rich in heavy minerals, glaucony and mica can give anomalous readings. in the correlation panels presented in this study (see plates 1–9), the borehole lithologies were described by the first author, with the following exceptions: fjand (dgu no. 76.635), fjelstervang (dgu no. 84.2649), lindved (dgu no. 116.1569), løgumkloster (dgu no. 159.739), ribe (dgu no. 140.42), rømø (dgu no. 148.52), tinglev (dgu no. 168.1378), uldum (dgu no. 1444), ulfborg (dgu no. 73.971),vester sottrup (dgu no. 169.799) and vollerup (dgu no. 160.1378). lithological descriptions of the latter boreholes are from the ‘jupiter’ well database at the geological survey of denmark and greenland (geus). all sample depths from boreholes are adjusted using the gamma-ray log in order to get true depths of the samples. thus there may be a disfig. 7. revised lithostratigraphic framework of the uppermost oligocene – miocene of onshore denmark, as presented here. r.: resen. plio.: pliocene. plio. zanclean messinian tortonian m åd e g ro u p r ib e g ro u p serravallian m io ce n e n eo ge n e 10 15 20 25 chattian langhian burdigalian aquitanian o lig o ce n e u p p er u p p er m id d le l o w er 5 p al ae o ge n e p er io d sw ne epochma age dinocyst zonation lithostratigraphy h. o. g. v. a. a. u. a. l. t. a. u. c. c. e. i. c. au. c. g. d. p. s. h. t. p. c. am. h. spp marbæk fm gram fm ørnhøj fm hodde fm arnum fm stauning mb odderup fm bastrup fm fasterholt mb vandel mb resen mb resen mb r. mb klintinghoved fm vejle fjord fm brejning fm brejning fm skansebakke mb kolding fjord mb øksenrade mb not included in this study billund fm hvidbjerg mb addit mb sydklint mb marine silt and clay marine sand fluvial sand and gravel hiatus brackish-water silt and clay coal bulletin 22_ gsb191-indhold 04/03/11 12.41 side 15 16 crepancy (usually less than 4 m) between depths indicated in the jupiter database (measured depth: md) and the depths assigned to the lithostratigraphic units in this study. the measured depth of cuttings samples is, however, indicated in the text. approximately 1000 km of 2d high-resolution seismic data have been used to correlate between boreholes and to investigate the overall architecture of the miocene succession. the correlations are also guided by dinocyst studies of most of the boreholes included here. these studies have resulted in a detailed dinocyst zonation (fig. 6; dybkjær & piasecki 2008, 2010). the geological age assigned to each lithostratigraphic unit is based primarily on this dinocyst stratigraphy (fig. 7). the lithostratigraphy of the uppermost oligocene – mio cene succession of onshore denmark is herein formally revised according to the guidelines presented by salvador (1994). nine lithostratigraphic units are revised and/or elevated in rank, 13 new lithostratigraphic units are erected. the oligocene to lowermost miocene brejning clay member, previously referred to the vejle fjord formation, is elevated to formation status; it includes the sydklint member and the øksenrade member. the miocene succession is subdivided into two groups, the ribe and måde groups. the ribe group consists of the vejle fjord, billund, klintinghoved, bastrup, arnum, and odderup formations. the vejle fjord formation includes the skansebakke member, the billund formation includes the hvidbjerg and addit members, the klintinghoved formation includes the kolding fjord member, the bastrup formation includes the resen member, the arnum formation includes the vandel member and the odderup formation includes the stauning and fasterholt members. the måde group comprises the hodde, ørnhøj, gram, and marbæk formations (fig. 7). it should be noted that particularly distinctive portions of individual formations are defined as members, but the formations are not subdivided at member level in their entirety. lithostratigraphic definition of units in complex interdigitating lithologies requires clear recognition of the lithological (or petrophysical in subsurface data) bounding criteria for formations and members. in this study, the following criteria were adopted. the sand-rich formations (e.g. billund, bastrup, odderup formations) possess over 75% sand and have a minimum thickness of 5 m; intercalated mudstone packets over 5 m thick are referred to the coeval marine, mud-rich formation (i.e. the vejle fjord formation in the case of the sand-rich billund formation). similarly, the marine, mud-dominated formations may contain subordinate sands; sand-rich intervals (with over 75% sand) that exceed 5 m in thickness are referred to the coeval sand formation. salvador (1994) and subsequent lithostratigraphic guidelines (nacsn 2005) discourage the use of stratigraphically alternating formations in interdigitating depositional systems; the practical disadvantages in outcropping terranes are clear. in subsurface lithostratigraphy, however, this practise is adopted on occasion (e.g. johnson & lott 1993) and is utilised here to emphasise the genetic integrity of the deltaic sandy systems. brejning formation new formation history.the brejning formation corresponds to the brejning clay member of the vejle fjord formation of larsen & dinesen (1959). name. after the town of brejning, south of vejle fjord (fig. 1). type and reference sections. the exposure at skansebakke, brejning (55°40´19.74´´n, 9°41´33.84´´e) forms the type section for the brejning formation (larsen & dinesen 1959, fig. 12). at low tide, the brejning formation is exposed in the basal, south-eastern part of the skansebakke profile at brejning. a borehole at brejning encountered a c. 4 m thick (–0.4 to –4.65 m) section referred to the brejning formation (larsen & dinesen 1959). the reference section is the outcrop at dykær, juelsminde (fig. 8). other exposures of the formation are found at sanatoriet and fakkegrav in the vejle fjord area, and at jensgård at the mouth of horsens fjord. in central jylland, the formation crops out at the sofienlund clay pit; in the limrevised lithostratigraphy bulletin 22_ gsb191-indhold 04/03/11 12.41 side 16 fjorden area, the formation is exposed at lyby and mogenstrup. periodically, the formation is exposed at søvind, sønder vissing, and in the ølst and hinge clay pits. the reference borehole section is the interval from 100.90 to 96.50 m (101–97 m md) in the andkær borehole (dgu no. 125.2017; fig. 8). thickness.the brejning formation is normally 2–4 m thick, but is over 20 m thick in a number of wells (plates 2, 5), and a 50 m thick succession referred to the brejning formation was encountered in the borg-1 borehole (plate 9). lithology. the brejning formation consists of greenish to brown, glaucony-rich clay with scattered pebbles (fig. 9). in the upper part, there is an increased content of organic matter, silt and sand. siderite concretions are also common in the upper part of the formation. the clay mine ralogy is dominated by illite, but smectite, kaolinite and gibbsite are also present (friis 1994; e.s. rasmussen 1995). mica is common in the upper part of the formation. log characteristics. high gamma-ray readings characterise the brejning formation (fig. 8); the lower part, in particular, may show extremely high gamma-ray values due to 17 fig. 8. reference sections for the brejning formation. the primary reference section is the dykær outcrop located south-west of juelsminde and the secondary reference section is the interval from 100.9 to 96.5 m in the andkær borehole. the accompanying legend is applicable to all outcrop and borehole logs shown in this study. vejle fj.: vejle fjord. 0 1 2 m cl si f mc p sand 100 101 95 m.b.s. gr 96 97 98 99 o lig o ce n e l o w er m io ce n e o lig o ce n e b re jn in g fm m io ce n e v ej le f j. fm b re jn in g fm e o ce n e sø vi n d f m v ej le f jo rd f m dykær outcrop andkær borehole dgu no. 125.2017 lithology clay/silt heterolith fine medium coarse gravel coal sedimentary structures erosional surface parallel bedding planar cross-bedding trough cross-bedding hummocky cross-bedding swaley cross-bedding inclined heterolithic bedding cross-lamination wavy bedding flaser bedding backflow ripples climbing ripples double clay layers concretion collapse structures biogenic structures bioturbation ophiomorpha rootlets fossils bivalve plant wood sponge spicules diatoms metres below surface gamma ray sand clast clast m.b.s. gr 40 cps 120 bulletin 22_ gsb191-indhold 04/03/11 12.41 side 17 18 the high content of glaucony (e.g. rødding borehole, plate 8), although expanded sections (e.g. borg borehole, plate 9) may show uniform intermediate values. fossils. the marine clay of the brejning formation contains a rich mollusc fauna (ravn 1907; eriksen 1937; schnetler & beyer 1987, 1990). marine microfossils, such as foraminifers (larsen & dinesen 1959; ulleberg 1987, 1994; laursen & kristoffersen 1999), calcareous nannofossils (von salis perch-nielsen 1994) and dinocysts (dybkjær 2004a, b; rasmussen & dybkjær 2005), are represented, and foraminifers and dinocysts are abundant and diverse. in the upper part of the formation, a gradual change/detoriation in the mollusc fauna was interpreted to reflect a shallowing-upward trend. similarly, in the dykær and jensgård exposures, the abundance and diversity of foraminifers (larsen & dinesen 1959) and dinocysts decrease in the upper part of the formation whereas the abundance and diversity of spores, pollen and freshwater algae increase (dybkjær 2004a, b; rasmussen & dybkjær 2005). echinoids, crinoids, asteroids, anthozoans, otoliths, sharks’ teeth, brachiopods, crustaceans and bryozoans have also been found. depositional environment. the brejning formation was deposited in a fully marine, sediment-starved environment (larsen & dinesen 1959; schnetler & beyer 1990; e.s. rasmussen 1995; rasmussen & dybkjær 2005). the water depth was probably more than 200 m in the norwegian–danish basin based on otoliths (schnetler & beyer 1990) and benthic foraminifera (c. morigi, personal communication 2009). the heights of clinoforms (offshore denmark) associated with early oligocene delta progradation indicate a minimum water depth of 200 m (danielsen et al. 1997), and since the late oligocene was warmer than the early oligocene (zachos et al. 2001), relatively deep water probably prevailed within the norwegian–danish basin during deposition of the brejning formation. schnetler & beyer (1990) reported a mixed mollusc fauna, some elements indicating deep marine conditions and some indicative of shallow water; the shallow marine fauna is most likely reworked, i.e. transported down the delta or shelf slope to the basin floor. on the ringkøbing–fyn high, shallower water prevailed. the upward increase in silt and sand indicates progradation of the shoreline in the latest oligocene associated with a relative sea-level fall (rasmussen & dybkjær 2005). boundaries. in southern and western jylland, the brejning formation rests with a sharp and erosional boundary on the eocene søvind marl formation (fig. 9; heilmannclausen et al. 1985). in this area, the boundary is marked by a distinct change in colour and grain size from the greenish grey clay of the søvind marl formation to the greenish brown and commonly silty brejning formation. the boundary may locally be intensively bioturbated and consequently more gradational. in central and northern jylland, the boundary is defined where dark brown clay of the branden formation (lower upper oligocene) is overlain by greenish glaucony-rich clay of the brejning formation. the base of the brejning formation is marked by a prominent shift to higher values on the gamma-ray log in the andkær borehole, but may locally be more gradational due to glaucony-filled burrows in the upper part of the søvind marl formation. the upper boundary is typically sharp and characterised by a change from greenish, dark brown, glaucony-rich clayey silt of the brejning formation to dark brown clayey silt of the overlying vejle fjord formation. a change in the degree of consolidation is also observed at the boundary in most parts of jylland from the well-consolidated sediments of the brejning formation to the relatively loose sediments of the vejle fjord formation. a gravel layer commonly occurs immediately above the upper boundary. at the type locality, the upper boundary is recognised by a disfig. 9. the brejning formation at øksenrade showing the lower part of the formation and the lower boundary with the underlying, light greenish-grey middle eocene søvind marl formation (photograph courtesy of peter warna-moors). red penknife for scale, c. 10 cm long. brejning fmbrejning fm søvind marl fmsøvind marl fm bulletin 22_ gsb191-indhold 04/03/11 12.41 side 18 tinct decrease in the content of glaucony passing from the brejning formation to the vejle fjord formation (larsen & dinesen 1959). the scattered glaucony grains found in the vejle fjord formation are reworked (e.s. rasmussen 1987). in central east jylland, the boundary is commonly characterised by a marked change from the sand deposits of the øksenrade member to the dark brown, clayey silt of the vejle fjord formation. 19 n brejning fm (oligocene) brejning fm (miocene) øksenrade mb sydklint mb vejle fjord fm skansebakke mb billund fm addit mb hvidbjerg mb klintinghoved fm kolding fjord mb bastrup fm resen mb arnum fm vandel mb odderup fm stauning mb fasteholt mb hodde fm, ørnhøj fm, gram fm marbæk fm a b d g e c f h 50 km fig. 10. the distribution of uppermost oligocene – miocene formations and members in denmark. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 19 20 distribution. the brejning formation is present in much of central and southern jylland but is typically absent on the ringkøbing–fyn high (fig. 10a). due to the diachronous nature of the upper boundary (see below), the youngest beds referred to the brejning formation are only present in southern and western jylland. the northern and eastern limit closely follows that of the miocene deposits (fig. 3). biostratigraphy. the deflandrea phosphoritica dinocyst zone of dybkjær & piasecki (2010) is recorded in the brejning formation. in addition, the chiropteridium galea zone is recorded in the upper part of the formation in the southern parts of jylland. geological age. the brejning formation is of late chattian to early aquitanian (latest late oligocene to earliest early miocene) age. the dinocyst stratigraphy indicates that the upper boundary of the brejning formation is diachroneous. in central parts of jylland, the boundary broadly correlates with the oligocene –miocene boundary (e.s. rasmussen 2004b; rasmussen & dybkjær 2005; dybkjær & rasmussen 2007). in the southern part of jylland, deposition of the glaucony-rich clay of the brejning formation apparently continued into the early aquitanian. subdivision. the brejning formation includes the sydklint member and the new øksenrade member. sydklint member history. a thin diatomite layer of oligocene age, unconformably overlying the lower eocene fur formation and overlain by upper oligocene micaceous clay, was observed in the cliff section at silstrup, near thisted, by heilmannclausen (1982). although noted by bøggild (1918), he apparently considered the layer to represent a glaciotectonically derived slice of the fur formation (see heilmannclausen 1997). the silicoflagellate assemblage in the diatomite layer was described by von salis (1993). the lithology of the oligocene diatomite layer and the contact to the underlying fur formation were described in more detail by heilmann-clausen (1997) who also proposed a model for the genesis of the diatomite. heilmann-clausen (1997) formally defined the unit as the sydklint member, which he provisionally referred to the vejle fjord formation. cl si f mc p sand sydklint outcrop 0 1 m u p p er o lig o ce n e e o ce n e b re jn in g fm fu r fm s. m b fig. 11. type section of the sydklint member. modified after heilmann-clausen (1997); for legend, see fig. 8, p. 17. s.: sydklint. fig. 12. contact between the light grey lower eocene fur for mation and the brown upper oligocene sydklint member at silstrup sydklint. a thin glauconitic layer occurs at the boundary between the two units. thalassinoides burrows extend from the glauconitic layer down into the topmost fur formation. knife for scale. glauconyglaucony fig. 13. photomicrograph of a vertical thin-section through the diatomaceous sydklint member. 100 μm bulletin 22_ gsb191-indhold 04/03/11 12.41 side 20 21 name. after the coastal cliff of silstrup sydklint, south of thisted (fig. 1). type section. the type section is the coastal cliff of silstrup sydklint (56°55´15.49´´n, 8°39´20.76´´e; fig. 11). thickness. the member is up to 28 cm thick. lithology. the sydklint member includes a basal 1–8 cm thick clay layer rich in coarse-grained glaucony and reworked clasts of the fur formation (figs 11, 12). sporadic extrabasinal pebbles and a single 25 cm large, partly glauconitised gneiss clast have been found in the basal layer. the glaucony-rich basal layer is succeeded by 20 cm of brown, clayey diatomite (fig. 13). log characteristics. the member is only recognised at outcrop, and log data are not available. fossils. the sydklint member contains well-preserved siliceous and organic-walled microfossils, including diatoms, silicoflagellates, sponge spicules, dinocysts, pollen and spores (fig. 13). depositional environment. the sydklint member was deposited in a marine, probably shelf environment. boundaries. the sydklint member has a sharp lower boundary separating the glaucony-rich basal layer from the underlying fur formation. an omission suite of shallow thalassinoides burrows extends 5–8 cm down into the topmost fur formation. the upper boundary is gradational over a few centimetres. distribution.the sydklint member is only known from outcrops at silstrup sydklint and nearby klovbakker at sundby, mors. biostratigraphy. the member is referred to the silicoflagellate distephanus speculum haliomma subzone of bukry (1981) by von salis (1993) and to the deflandrea phosphoritica dinocyst zone of dybkjær & piasecki (2010) by heilmann-clausen (c. heilmann-clausen, personal communication 2010). geological age. the sydklint member is of late chattian (latest late oligocene) age. øksenrade member new member history. the succession defined here as the øksenrade member was termed ‘middelfart malm’ by l.b. rasmussen (1975). equivalent oolitic ironstones cropping out at jensgård at the mouth of horsens fjord were described by friis et al. (1998). name. after øksenrade skov, just north of the coastal type locality (fig. 1). type and reference sections. the type section is the coastal cliff facing fænø sund, south of øksenrade skov, middelfart (55°29´39.61´´n, 9°42´47.29´´e; fig. 14). the reference section is the interval from 212 to 210 m (214–212 m md) in the borehole at gadbjerg (fig. 14; dgu no. 115.1474). fig. 14. type and reference sections of the øksenrade member (ø.). the type section is the øksenrade outcrop located south-west of middelfart; the top of the member is not seen. the reference section is the interval from 212 to 210 m in the gadbjerg borehole; for legend, see fig. 8, p. 17. øksenrade outcrop 0 1 2 3 4 m cl si f mc p sand 230 220 210 200 m.b.s. gr gadbjerg borehole dgu no. 115.1474 l o w er m io ce n e u p p er o lig o ce n e u p p er o lig o ce n e b re jn in g fm ø ks en ra d e m b e o ce n e v ej le f jo rd f m ø . b re jn in g fm sø vi n d m ar l fm 40 cps 100 bulletin 22_ gsb191-indhold 04/03/11 12.41 side 21 22 thickness. the member is c. 1 m thick at the type locality (fig. 14), but the top is not seen; boreholes indicate a maximum thickness of 5 m (fig. 14; plate 8). lithology. the øksenrade member is composed of reddish ooids and grey, well-sorted, fine-grained quartz sand (figs 15, 16) and ranges lithologically from a sand with dispersed ooids to a sandy ooid grainstone. at jensgård, the øksenrade member consists of planar cross-bedded sand, sets are up to 40 cm thick and typically show asymptotic toesets. the foresets are inclined towards the north. the ooids are composed of concentric layers of goethite, commonly with a core of glaucony grains or pellets (h. friis, personal communication 2010); at the type section, shells or quartz grains also form ooid cores. the cement consists of siderite with some calcite; the iron content of the sediment is up to 30% (e.s. rasmussen 1987). moulds of mollusc shells are common. log characteristics. the øksenrade member is typified by relatively low gamma-ray readings (fig. 14) but distinct spikes may occur due to horizons rich in glaucony. fossils. the øksenrade member is characterised by abundant moulds of mollusc shells (l.b. rasmussen 1975; gravesen 1990). depositional environment. the øksenrade member was deposited above storm wave base as indicated by cross-bedding (rasmussen & dybkjær 2005). the bivalve and gastropod faunas (l.b. rasmussen 1975) also indicate a shallow-water depositional environment. the transgressive lag that is locally found on the ringkøbing–fyn high at the base of the vejle fjord formation is indicative of exposure and terrestrial sedimentation prior to transgressive reworking (rasmussen & dybkjær 2005). such a shallowing and local emergence at the transition from the oligocene to the miocene is also indicated by the presence of freshwater algae in the upper part of the brejning formation (rasmussen & dybkjær 2005). boundaries. the øksenrade member rests with a sharp erosional boundary on the undifferentiated brejning formation beneath (figs 15, 16). the lower boundary is also marked by a change from dark brown, clayey silt with scatfig. 15. brejning formation and the øksenrade member in the coastal cliff at øksenrade skov, south-west of middelfart. this outcrop constitutes the type section of the øksenrade member; spade for scale (c. 1.3 m long). øksenrade mb brejning fm øksenrade mb brejning fm bulletin 22_ gsb191-indhold 04/03/11 12.41 side 22 23 tered sand lenses to fine-grained, reddish sand; on the gamma-ray log, this facies shift is reflected by a shift towards lower values. the upper boundary is characterised by a marked change from the sand deposits of the øksenrade member to dark brown, clayey silts of the vejle fjord formation; this boundary is marked by a prominent shift on the gamma-ray log from low to high values. distribution. the øksenrade member is present in east jylland and the extreme western part of fyn, from horsens in the north to middelfart in the south (fig. 10a). the westernmost limit is defined by exposures at gadbjerg near give where the member occurs on a footwall crest at the boundary fault of the brande trough. biostratigraphy. no samples from this member have been analysed for palynology; the mollusc fauna is non-specific. geological age. based on stratigraphic context, the øksenrade member is considered of latest chattian (latest late oligocene) age. fig. 16. boundary between the brejning formation and the øksen rade member, marked by a distinct colour change from dark brown clayey silt to red sand. lens cap for scale. øksenrade mb brejning fm øksenrade mb brejning fm bulletin 22_ gsb191-indhold 04/03/11 12.41 side 23 24 history. non-fossiliferous sand and gravel encountered below 125.6 m in a borehole at ribe were defined as the ribe formation by sorgenfrei (1958). the borehole terminated at a depth of 127 m and thus the base of the formation was never defined. l.b. rasmussen (1961) suggested that a succession of quartz gravel and sand with some lignite between 255.7 and 144.5 m in the arnum-1 borehole should be referred to the ribe formation. he further indicated that the fluvio-deltaic, brown-coal-bearing succession around silkeborg and skanderborg may be correlative with the ribe formation. in this stratigraphic revision, however, the fluvio-deltaic deposits at silkeborg are referred to the vejle fjord and billund formations. the ribe formation was included in the stratigraphic chart of l.b. rasmussen (1961) where it was suggested to encompass the fluvio-deltaic deposits below the odderup formation. the age of the formation was indicated as early to early middle miocene (fig. 4). during the last decade, detailed biostratigraphic and sequence stratigraphic studies of the lower miocene succession have been carried out (e.s. rasmussen 2004b; dybkjær 2004a; rasmussen & dybkjær 2005; e.s. rasmussen et al. 2006; dybkjær & piasecki 2010). these studies have revealed that the stratigraphy of the lower miocene deposits is more complicated than formerly believed. the fluvio-deltaic sediments that are so characteristic of the lower miocene – lower middle miocene succession are thus here defined as the ribe group. the introduction of the bastrup formation, which replaces the ribe formation in southern jylland, is also consistent with the new lithostratigraphy of schleswig-holstein, northern germany (rasser et al. 2008; knox et al. 2010). here the bastrup formation was adopted to represent lower miocene fluvio-deltaic sands of burdigalian age, based on a study of the kasseburg cored borehole near hamburg (k. gürs, personal communication 2006; rasser et al. 2008; knox et al. 2010). the ribe group correlates with the upper part of the hordaland group as applied in the north sea region, ribe group new group store vorslunde borehole dgu no. 104.2325 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 m.b.s. gr r ib e g ro u p v ej le f jo rd f m b ill u n d f m k lin ti n gh o ve d f m b as tr u p f m a rn u m f m o d d er u p f m b r. fm l o w er m io ce n e o lig o . 0 cps 80 fig. 17. the full development of the ribe group is illustrated by the interval from 219 to 1 m in the store vorslunde borehole, north-east of vejle; for legend, see fig. 8, p. 17. br.: brande. oligo.: oligocene. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 24 25 including offshore denmark in the norwegian–danish basin (deegan & scull 1977; hardt et al. 1989; schiøler et al. 2007). name. after the town of ribe (fig. 1). type area. the type area of the ribe group is central and east jylland. in the gravel pit at voervadsbro in central jylland (fig. 1), both marine sand and fluvial sand and gravel of the ribe group are exposed. in the store vorslunde borehole (fig. 17; dgu no. 104.2325) a complete section through the group is represented in the interval from 219 to 1 m (220–1 m md). the group crops out at klintinghoved in southern jylland, at rønshoved, hagenør, børup, hindsgavl, galsklint, hvidbjerg, brejning, sanatoriet, fakkegrav, dykær and jensgård in eastern jylland, at addit, salten, isenvad and abildå in central jylland and at gyldendal, søndbjerg, lyby, skyum bjerge, skanderup and lodbjerg in the limfjorden area. thickness. the group is 218 m thick in the store vorslunde borehole. a thickness of c. 200 m is common in the norwegian–danish basin and in most places on the ringkøbing–fyn high. in the tinglev borehole, located in the tønder graben, more than 200 m has been penetrated without reaching the lower boundary of the group (plate 1). reduced thicknesses are seen in the eastern part of jylland, partly due to erosion during the pleistocene. lithology. the group consists of three cycles of alternating mud-rich and sand-rich units with some intercalation of coal beds, especially in the upper cycle (odderup formation); each cycle, 50 to 100 m thick, represents a coarsening-upward cycle. the sands are typically mediumto coarse-grained, quartz-rich with a variable mica content. various types of cross-bedding, including tabular, trough, hummocky and swaley cross-stratification, characterise the sand-rich units. the sand grains are normally subto wellrounded. well-rounded pebbles of quartz, quartzite and fig. 18. quartz-rich sand and pebbles from the miocene fluvio-deltaic deposits; photographs courtesy of p. warna-moors. a: pebbles of quartz, quartzite and chert; these are commonly found near sequence boundaries, associated with transgressive lags or within fluvial channels. note that the clasts are up to 3 cm in diameter. b: granules and coarse-grained sand of the billund formation. c: fineand medium-grained sand of the billund formation. a b c bulletin 22_ gsb191-indhold 04/03/11 12.41 side 25 26 chert up to 4 cm in size (fig. 18) commonly occur in the upper part of the units near sequence boundaries (in transgressive lags or fluvial channels). fossils occur only sporadically in the sand-rich units. the micaceous, mud-rich portions of the group are typically homogeneous, with some intercalation of laminated mud intervals as well as discrete sand layers. the sand layers are commonly hummocky cross-stratified or represent tidal rhythmites. the clay mineral association is dominated by illite, kaolinite and gibbsite (fig. 19); pyrite is a very common authigenic mineral. the coal beds are found associated with cross-stratified fluvial sands and muds, and also cap shoreface/beach sands and lagoonal muds. the coal beds are limited to the norwegian–danish basin where they typically reach thicknesses of 2–3 m; the thickest succession has been recorded in the fasterholt area, where there is a cumulative thickness of about 9 m of coal. lithological details of the respective formations of the ribe group are given below under the individual formation descriptions. log characteristics. the typical log pattern shows three cycles of decreasing-upward gamma-ray values (fig. 17). the gamma-ray log is generally characterised by a serrated pattern, but distinct gamma-ray spikes are common in the lower part of each cycle; in the upper cycle (the arnum and odderup formations), high gamma-ray spikes occur throughout the succession. in the northern part and also locally in the southern part, decreasing gamma-ray values are commonly observed in the upper part of each cycle. for more detailed descriptions, see the individual units below. fossils. molluscs occur abundantly in the marine and nearshore deposits and plant fossils are locally abundant in the terrestrial deposits. more detailed descriptions of the fauna/flora are given below in the definitions of the formations and members. depositional environment. the ribe group was deposited by delta systems prograding from the north and north-east towards the south and south-west. deposition of the first cycle (billund formation) was strongly controlled by the topography formed during early miocene inversion tectonism (rasmussen & dybkjær 2005; hansen & rasmussen 2008; e.s. rasmussen 2009a). during the deposition of this cycle, the so-called ringkøbing and brande lobes were focussed particularly within structural lows, the brande trough and the rødding graben (hansen & rasmussen 2008). east of the main delta lobes, spit and barrier-complexes developed due to shore-parallel transport of sand that was delivered from the river mouths of the delta systems (rasmussen & dybkjær 2005; hansen & rasmusen 2008). fluvial sands interpreted as braided river system deposits (hansen 1985; jesse 1995; e.s. rasmussen et al. 2006) dominate in the northern part. the second cycle (bastrup formation) shows a more evenly distributed progradational pattern across jylland. due to the lack of outcrops of this part of the miocene succession, detailed sedimentology has not been carried out. judging from borehole data, there are no indications of widespread spit and barrier complexes. as for the first cycle, fluvial systems dominate the upper part of the succession. log and seismic data (e.s. rasmussen et al. 2007; e.s. rasmussen 2009b) indicate that a meandering fluvial sysfig. 19. x-ray diffractogram of the clay fraction from muds of the vejle fjord formation (ribe group). note that gibbsite is present indicating that the source area was heavily weathered. full line = untreated samples, dotted line = glycolated samples and dashed line = samples heated to 500°c. modified from e.s. rasmussen (1995). da: lattice separation (in angstrom). 5 20°1020 15 5 73 10 18 da (0 0 2 ) g ib b si te (0 0 1 ) g ib b si te (0 0 1 ) k ao lin it e (0 0 2 ) ill it e (0 0 1 ) ill it e (0 0 1 ) sm ec ti te bulletin 22_ gsb191-indhold 04/03/11 12.41 side 26 27 tem was widespread, although local or periodic development of braided fluvial systems may have taken place. the third and final cycle (the odderup formation) was deposited in a prograding coastal plain with widespread coal formation within the norwegian–danish basin, whereas clean fluvial sand dominates the ringkøbing–fyn high area. boundaries. the lower boundary is commonly sharp, being defined where greenish to brownish, glaucony-rich clay and silt is overlain by dark brown, organic-rich mud. over much of jylland, the boundary is also marked by a change in the degree of consolidation, from the well-consolidated sediments of the oligocene brejning and branden formations to poorly consolidated ribe group sediments. the boundary may be marked by a gravel lag or sand bed. due to intense bioturbation, the boundary may be locally blurred. in central east jylland, the boundary is characterised by a marked change from the sand deposits of the øksenrade member to dark brown clayey silt of the vejle fjord formation of the ribe group. the upper boundary is sharp, being marked by a thin gravel layer that separates the white, fine-grained sand of the uppermost ribe group (odderup formation) from the dark brown mud of the succeeding måde group. this is reflected by a prominent shift on the gamma-ray log towards high gamma-ray values. distribution.the ribe group is present over most of jylland. the northern and eastern limits of the group closely follow the lower boundary of the miocene deposits (fig. 4) geological age. the ribe group is of aquitanian – early langhian (early miocene – earliest middle miocene) age. subdivision. the ribe group is divided into six formations: the aquitanian vejle fjord and billund formations, the uppermost aquitanian – lower burdigalian klin tinghoved and bastrup formations and the upper burdigalian – lower langhian arnum and odderup for mations (fig. 7). vejle fjord formation redefined formation general. the marine, clay-dominated vejle fjord formation interdigitates north-eastwards with the fluvio-deltaic, sandrich billund formation. these two formations thus alternate up-section in some boreholes (e.g. plates 2–8). fig. 20. type section of the vejle fjord formation and the skansebakke member at skansebakke, brejning; for legend, see fig. 8, p. 17. 0 1 2 3 4 5 6 7 8 m cl si f mc p sand skansebakke outcrop v ej le f jo rd f m sk an se b ak ke m b l o w er m io ce n e bulletin 22_ gsb191-indhold 04/03/11 12.41 side 27 28 dykær outcrop 0 1 2 3 4 5 6 7 8 9 10 m 10 11 12 13 14 15 16 cl si f mc p sand cl si f mc p sand b ill u n d f m l o w er m io ce n e h vi d b je rg m b b ill u n d f m v ej le f jo rd f m b re jn in g fm l o w er m io ce n e o lig o ce n e sk an se b ak ke m b h vi d b je rg m b fig. 21. primary reference section of the vejle fjord formation at dykær, south-west of juelsminde; for legend, see fig. 8, p. 17. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 28 29 history. the vejle fjord formation was defined by larsen & dinesen (1959). the formation was originally defined as the succession from the base of the brejning clay member to the top of the vejle fjord sand member. for stratigraphic and practical reasons, the brejning clay member is herein removed from the vejle fjord formation and elevated to the status of formation (see above); redefinition of the vejle fjord formation is therefore necessary. revision is also needed because of the large amount of data acquired during the last decade, which has shed new light on the depositional system (dybkjær & rasmussen 2000; rasmussen & dybkjær 2005). sediments referred by christensen & ulleberg (1973) to the upper sofienlund formation are assigned here to the vejle fjord formation; the sofienlund formation is abandoned. name. after vejle fjord in east jylland (fig. 1). type and reference sections. the type section is the skansebakke outcrop at brejning 55°40´19.74´´n, 9°41´33.84´´e; figs 1, 20). the outcrop reference section is defined at dykær near juelsminde (figs 1, 21). other exposures in the vejle fjord area are brejning hoved, sanatoriet, fakkegrav and jensgård. it is further exposed at hindsgavl near middelfart, and the formation crops out at skyum bjerge, lyby, mogenstrup and skanderup (mors) in the limfjorden area (fig. 1). the secondary reference section is the store vorslunde borehole (dgu no. 104.2325) (fig. 22), in the interval from 219 to 160 m (220–161 m md). thickness. the formation is c. 20 m thick at the type locality though neither the base nor the top are seen; the formation is about 18 m thick in the nearby andkær borehole (see plate 1). in the western part of jylland, it may reach a thickness of up to c. 100 m, as exemplified by the holstebro borehole (plate 4). lithology. the vejle fjord formation consists mainly of dark brown clayey silt (fig. 23). in some areas, it is dominated by laminated, greenish-grey sand and dark brown, clayey silt. sand stringers up to a few centimetres thick may occur. locally, the formation is composed of waveinfluenced heterolithic mud and sand showing hum mocky cross-stratification (figs 24, 25); the heterolithic succession is commonly characterised by double clay layers and climbing ripples. soft-sediment deformation structures occur locally. trace fossils occur in places in the vejle fjord formation. log characteristics. the formation is characterised by intermediate gamma-ray values (fig. 22). the log pattern is serrated and shows both decreasingand increasing-upward trends throughout the succession. fossils. the vejle fjord formation contains an impoverished mollusc fauna (ravn 1907; eriksen 1937; schnetler & fig. 22. secondary reference section of the vejle fjord formation: the interval from 219 to 160 m in the store vorslunde borehole, north-east of vejle; for legend, see fig. 8, p. 17. store vorslunde borehole dgu no. 104.2325 220 210 200 190 180 170 160 150 140 m.b.s. gr v ej le f jo rd f m b ill u n d f m b ra n d e fm l o w er m io ce n e o lig o ce n e 0 cps 80 bulletin 22_ gsb191-indhold 04/03/11 12.41 side 29 30 beyer 1987, 1990). the foraminifer fauna (larsen & dinesen 1959; laursen & kristoffersen 1999) and the dinocyst flora (dybkjær 2004 a, b; rasmussen & dybkjær 2005) are similarly impoverished within this formation, although the abundance of dinocysts is locally very high, albeit restricted to a few species. depositional environment. the vejle fjord formation was deposited in a brackish to fully marine depositional environment. brackish-water conditions predominated within the norwegian–danish basin in the early phase of deposition as a consequence of the elevated ringkøbing–fyn high (rasmussen & dybkjær 2005; e.s. rasmussen 2009a). as sea level rose during the early miocene, fully marine conditions were re-established and the water depth was c. 100 m in the norwegian–danish basin and probably less than 30 m on the ringkøbing–fyn high. most of the vejle fjord formation was deposited in a prodelta environment. the thickest developments of the formation are associated with inter-lobe depositional environments. boundaries. the lower boundary is typically sharp, being characterised by a change from greenish dark brown, glaucony-rich, clayey silt to dark brown, clayey silt. a change in the degree of consolidation is observed at the boundary over much of jylland, relatively loose sediments of the vejle fjord formation overlying well-consolidated sediments of the brejning formation. a gravel layer is commonly found at the lower boundary. at the type locality, the lower boundary is marked by a distinct decrease in the content of glaucony (larsen & dinesen 1959); the scattered glaucony grains found in the vejle fjord formation are reworked (e.s. rasmussen 1987). in central east jylland, the boundary is commonly characterised by a marked change from the sand deposits of the øksenrade member (upper brejning formation) to dark brown, clayey silt of the vejle fjord formation. recognition of the lower boundary of the formation in subsurface data is based on both lithological and petrophysical data. the gamma-ray response at the boundary is variable. where the upper brejning formation is mud-rich, as in the reference section for the formation (fig. 8), gamma-ray values fall at the boundary to intermediate levels. in contrast, where the upper brejning formation is sand-rich (e.g. fig. 14), or where the vejle fjord formation overlies deltaic sands referred to the billund formation (e.g. stakroge and assing mølleby boreholes, plate 3), the fig. 23. type section of the vejle fjord formation at skansebakke, brejning; spade for scale (c. 1.3 m long). bulletin 22_ gsb191-indhold 04/03/11 12.41 side 30 31 lower boundary is defined by an abrupt increase in gammaray values. although in a number of wells the boundary can be difficult to position based on log data alone, lithological evidence (e.g. the presence of a gravel layer, glaucony content, clay colour and consolidation) can aid identification (e.g. resen and mausvig boreholes, plate 5). the upper boundary is typically defined where clayey, organic-rich silty sediments of the vejle fjord formation are succeeded by sand-rich deposits (> 75% sand) with a minimum thickness of 5 m; the overlying sands are referred either to the billund formation or to the kolding fjord member of the klintinghoved formation (e.g. plate 6). on the gamma-ray log, this boundary may show a marked decrease in gamma-ray readings where overlain by a discrete sand unit, or a gradual but steady decrease in gammaray readings reflecting a transitional, interbedded, sandrich unit at the base of the overlying formation. where the billund formation is absent in south and west jylland, the vejle fjord formation is succeeded by the clay-rich klintinghoved formation, and the boundary can be difficult to locate in detail. in the rødding and føvling boreholes (plate 8), for example, the two formations are lithologically very similar although the clayey silts of the vejle fjord formation may be slightly more consolidated. the boundary can typically be picked on the gamma-ray log, however, at a minor or moderate upward increase in values, commonly capping a weak coarsening-upward succession (decreasing-upward gamma-ray values). fig. 24. hummocky cross-stratified sand in the upper part of the vejle fjord formation at jensgård, east of horsens (fig. 1); the dipping, weakly deformed attitude of these strata is due to glaciotectonics. fig 25. hummocky cross-stratified sand with burrows (scolicia isp.) from the vejle fjord formation at skyum (fig. 1). note that the sand layer is only burrowed in the upper part. most of the vejle fjord formation was deposited as alternating sand and clayey, silt layers, but due to bioturbation any stratification was later destroyed and only the thicker storm sand layers were preserved. knife blade for scale. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 31 distribution. the formation is present over much of jylland with the exception of the southern and westernmost parts (fig. 10b). the northern and eastern limit closely follows the overall outcrop pattern of the miocene deposits (fig. 4). biostratigraphy. the chiropteridium galea and the homotryblium spp. dinocyst zones of dybkjær & piasecki (2010) are recorded in the vejle fjord formation. geological age. the vejle fjord formation is of aquitanian (earliest early miocene) age. subdivision. the vejle fjord formation includes the skansebakke member. skansebakke member revised member history. sediments referred here to the skansebakke member were formerly assigned to the vejle fjord sand member by larsen & dinesen (1959); the member is renamed here in accordance with modern lithostratigraphic guidelines. name. after the outcrop of the type section at skansebakke, brejning, on the south coast of vejle fjord. type and reference sections. the type section is the outcrop at skansebakke (55°40´19.74´´n, 9°41´33.84´´e; fig. 1). it is also exposed at brejning hoved, sanatoriet, fakkegrav and dykær. the reference section is the interval from 91.10 to 79 m (92–79 m md) in the andkær borehole (dgu no. 125.2017; fig. 26). thickness. at the type locality, the member is c. 7 m thick (top not seen); the member is c. 12 m thick at brejning hoved and 7 m thick at sanatoriet. lithology. the skansebakke member consists of alternating layers of fine-grained, well-sorted, yellowish sand and brownish clay (fig. 27). the sand beds are sharp-based and homogenous to evenly laminated. the sand beds are commonly capped by waveand current-ripples. the trace fossils arenicolites isp. and macaronichnus isp. are common, and ophiomorpha isp. is sporadically distributed (friis et al. 1998). the pyrite content is relatively high compared to the overlying billund formation, resulting in the yellowish colour in exposed sections (m. olivarius, personal communication 2010). log characteristics. the member is characterised by low gamma-ray readings with a serrated pattern (fig. 26), reflecting the alternation of sand and mud beds. fossils. the skansebakke member contains an impoverished mollusc fauna (ravn 1907; eriksen 1937). the foraminifer fauna (larsen & dinesen 1959) and the dinocyst flora (dybkjær 2004 a, b; rasmussen & dybkjær 2005) are also impoverished within this member. depositional environment. the skansebakke member is interpreted as having been deposited in a lagoonal depositional environment (larsen & dinesen 1959; friis et al. 1998; rasmussen & dybkjær 2005). the sand beds were deposited as washover fans on a backbarrier flat during the main degradation of minor spit and barrier systems formed along elevated parts of the ringkøbing–fyn high. boundaries. the lower boundary is placed at the base of the first significant sand layer separating dark brown, clayey silt from a succession dominated by interbedded yellowish fine-grained sand and dark brown to brown, silty clay. on the gamma-ray log, the lower boundary is placed at a minor, but distinct decrease in gamma-ray readings. the upper boundary is defined by the distinct change from yellowish, fine-grained sand to white, fineto medium-grained 32 andkær borehole dgu no. 125.2017 100 90 80 70 m.b.s. gr v ej le f jo rd f m l o w er m io ce n e o li. b j. fm b ill u n d f m h vi d b je rg m b sk an se b ak ke m b 20 cps 100 fig. 26. the reference section of the skansebakke member is the interval from 91.10 to 79 m in the andkær borehole. note that the sand-rich nature of the interval indicated by the gamma-ray log is not reflected by the lithological sample data; for legend, see fig. 8, p. 17. bj.: brejning. oli.: oligocene. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 32 33 sand of the hvidbjerg member (billund formation). this boundary is only documented in the andkær borehole where the gamma-ray log changes from serrated, low–intermediate gamma-ray readings of the skansebakke member to more consistently low gamma-ray values of the hvidbjerg member. distribution. the skansebakke member is restricted to central east jylland and is exposed along the coast of vejle fjord (fig. 10b). in the subsurface, this member is only recognised in the andkær borehole. biostratigraphy. the chiropteridium galea and the homotryblium spp. dinocyst zones of dybkjær & piasecki (2010) are recorded in the skansebakke member. geological age. the skansebakke member is of aquitanian (earliest early miocene) age. fig. 27. alternating fine-grained sand and clay of the skansebakke member at skansebakke. the clay was deposited in a lagoon and the sand was deposited as washover fans during the degradation of a barrier island associated with an early miocene transgression. spade for scale (c. 1.3 m long). bulletin 22_ gsb191-indhold 04/03/11 12.41 side 33 34 billund formation new formation name. after the town of billund (fig. 1). type and reference sections. the type section of the billund formation is the interval from 235 m to 184 m (235–185 m md) in the billund borehole (dgu no. 114.1857, 55°43´08.53´´n, 9°08´33.98´´e; fig. 28). the reference section is the interval from 160 to 126 m (161–128 m md) in the store vorslunde borehole (dgu no. 104.2325; fig. 28). thickness. in the type section, the formation is 51 m thick; the maximum thickness of 77 m has been found in the hammerum borehole (plate 6). lithology. the billund formation is primarily known from the subsurface but is exposed at a number of localities. in the lillebælt area, the formation is exposed at børup, galsklint, hindsgavl, røjle and rønshoved and in the vejle fjord region at dykær, fakkegrav and hvidbjerg. in central jylland, the formation can be observed at addit, salten and voervadsbro, and at søndbjerg and lyby in northern jylland. it is composed of fineto coarse-grained sand with some gravel or pebble-rich beds (fig. 29). the formation consists of almost pure quartz sand and includes clasts of quartzitic sandstone with subordinate mica and heavy minerals. clasts of well-rounded chert occur locally. pebbly horisons are common in the upper part and at the base of fluvial channels; clasts up to 4 cm occur in erosional scours within steep clinoform units. the formation is characterised by both coarsening-upward and fining-upward depositional patterns. fine-grained sand units which are commonly hummocky cross-stratified, occur in the lower part of the formation and in eastern sections. the upper part is commonly dominated by swaley cross-stratified sand or sharp-based sand with a homogeneous or laminated lower part capped by wave ripples. the trace fossils ophiomorpha isp. and skolithos isp. are common (fig. 29; friis et al. 1998; rasmussen & dybkjær 2005). in the northern area, the formation is dominantly composed of cross-bedded sand with a range of set thicknesses. soft sediment deformation structures are commonly seen. some sections show an interval of interbedded, fine-grained, wave-rippled sands, muds and coals, sandwiched between two sand bodies with an overall sheet geometry. root horizons and tree stumps are locally present (weibel 1996; e.s. rasmussen et al. 2007). in the eastern area, where the formation crops out, the sands are characterised by hummocky and swaley cross-stratification and homogeneous to laminated sand beds commonly capped by wave ripples; tidal bundles are also present (fig. 30). the interbedded muds and heteroliths are dark brown in the northern part due to a high content of organic matter. in the southern area, the mud is light brown and typically thinner bedded, occurring interbedded with storm sand beds. log characteristics. the formation is generally characterised by low gamma-ray values. in some boreholes, the lower part is characterised by a serrated lower part with generfig. 28. type and reference sections of the billund formation. the type section is the interval from 235 to 184 m in the billund borehole and the reference section is the interval from 160 to 126 m in the store vorslunde borehole; for legend, see fig. 8, p. 17. bra.: brande. olig.: oligocene. billund borehole dgu no. 114.1857 store vorslunde borehole dgu no. 104.2325 l o w er m io ce n e o lig o . v ej le f jo rd f m b ra . f m k lin ti n gh o ve d f m b ill u n d f m l o w er m io ce n e e o ce n e v ej le f jo rd f m b ill u n d f m sø vi n d f m 240 230 220 210 200 190 180 170 160 220 210 200 190 180 170 160 150 140 130 120 110 100 m.b.s. m.b.s. grgr 20 cps cps 100 0 80 bulletin 22_ gsb191-indhold 04/03/11 12.41 side 34 35 ally higher gamma-ray values (e.g. hammerum borehole, plate 6). in the type borehole, the billund formation shows consistently low gamma-ray readings (fig. 28). fossils. the billund formation contains fossil wood (weibel 1996), leaves and seeds (ravn 1907) but also marine molluscs (e.g. in the ‘brøndum blokke’; friis 1995). foraminifers and dinocysts are present locally (laursen & kristoffersen 1999; e.s. rasmussen et al. 2006). depositional environment. the billund formation was deposited as a delta system prograding from the north and north-east towards the south and south-east. the well-constrained palaeogeographical setting is based on high-resolution seismic data and facies distribution (rasmussen & dybkjær 2005; hansen & rasmussen 2008). progradation took place in association with an early miocene inversion phase (e.s. rasmussen 2009a), and the distribution of the delta lobes was consequently strongly controlled by the antecedent topography. two major lobes, the ringkøbing and brande lobes, were mapped by hansen & rasmussen (2008). the billund delta complex was deposited as wavedominated deltas (rasmussen & dybkjær 2005; hansen & rasmussen 2008; e.s. rasmussen 2009b). the southeastward longshore currents that prevailed during the early miocene resulted in deposition of spit and barrier complexes south-east of the main delta lobes (hvidbjerg member). the most coarse-grained part was deposited in steeply dipping clinoformal packages deposited during falling sea-level (hansen & rasmussen 2008; e.s. rasmussen 2009b) and within incised valleys (addit member). boundaries. the lower boundary is defined by a change from clayey, organic-rich silty sediments of the vejle fjord formation to sand-rich deposits; as noted earlier, recognition of the billund formation requires a minimum sand thickness of 5 m and a sand–mud ratio of over 75%. locally, for example in the type section at billund, sand referred to the billund formation overlies the eocene søvind marl; in such sections, the base of the formation is a significant fig. 29. marine sand and fluvial gravel and sand of the billund formation exposed at voervadsbro. note the skolithos burrows (centre) indicating a marine depositional environment. the lower boundary of the fluvial deposits (addit member) is at the base of the gravel layer (dashed line). the illustrated section is 2 m high. addit mb billund fm addit mb billund fm bulletin 22_ gsb191-indhold 04/03/11 12.41 side 35 36 hiatal surface. on the gamma-ray log, the lower boundary is identified by a marked decrease in gamma-ray readings. in some sections (e.g. store vorslunde, fasterholt boreholes, plate 2), the shift from mudto sand-rich deposits is gradational and reflected by a gradual but steady decrease in gamma-ray readings; the boundary is placed according to the criteria described above. the upper boundary is placed at the change from sandrich deposits of the billund formation to the predominantly dark brown, silty clays of the klintinghoved formation or the vejle fjord formation. at outcrop, the boundary is often erosive and overlain by a gravel lag or sand layer showing a fining-upward trend; the base of the gravel lag or sand layer forms the upper boundary. on the gamma-ray log, the upper boundary shows a variety of motifs. in boreholes where the sandy billund formation is succeeded by mud-dominated facies of the klintinghoved formation, the boundary is defined at an abrupt increase in values. where the lower klintinghoved formation includes gravel and sand layers succeeded by mud-rich facies (e.g. egtved borehole, plate 7), the boundary is placed at the base of a prominent shift to lower gamma-ray values that is succeeded by a general upward increase in values. where gamma-ray readings are strongly serrated, the boundary is placed at the base of the most coarse-grained sand or gravel layer found in the lithological descriptions. distribution. the billund formation is distributed in central jylland (fig. 10c). although beyond the formal boundaries of the formation, a sand-rich succession reported from the subsurface of the north sea may represent the westernmost lobe of the billund delta complex (hansen & rasmussen 2008). biostratigraphy. the chiropteridium galea and the homotryblium spp. dinocyst zones of dybkjær & piasecki (2010) are recorded in the billund formation. geological age. the billund formation is of aquitanian (earliest early miocene) age. subdivision. the billund formation includes two members: the hvidbjerg member and the addit member. hvidbjerg member new member general. the new hvidbjerg member represents a particular facies variant of the billund formation, dominated by spit deposits. the diagnostic features are only convincingly recognised at outcrop; the member is thus only recognised in the vejle fjord area in outcrops and closely adjacent boreholes. the member also crops out at søndbjerg and lyby in the limfjorden area (fig. 1). history. the succession of white sands at hvidbjerg strand was studied by larsen & dinesen (1959); these authors refrained from including the ‘hvidbjerg sand’ in the vejle fjord formation due to contrasting heavy mineral suites in these two units. name. after the outcrop at hvidbjerg strand on the south coast of vejle fjord (fig. 1). type and reference sections. the type section of the hvidbjerg member is the coastal exposure at hvidbjerg strand on the south coast of vejle fjord (55°38´24.58´´n, 9°44´39.22´´e; figs 31, 33). other exposures are at sanatoriet, fakkegrav and dykær in the vejle fjord area, at pjedsted north-west of fredericia and at hindsgavl, galsklint, børup and rønshoved in the lillebælt area. the sand crops out at two localities in the limfjorden area, at søndbjerg and lyby. the reference section is the interval bulletin 22_ gsb191-indhold 04/03/11 12.41 side 36 37 from 79 to 58 m (79–58 m md) in the andkær borehole (dgu no. 125.2017; fig. 32). thickness. the member is 28 m thick in the type section at hvidbjerg (fig. 31). in the outcrops of the lillebælt area, it can attain 13 m but is rarely thicker than 6 m. in the subsurface, the member is recognised in the andkær borehole (reference section, 21 m thick) and the lillebælt borehole (c. 11 m). lithology. the hvidbjerg member consists of white, fineto medium-grained sand with a few pebble layers (fig. 33). the sand beds are dominated by sharp-based, structureless to evenly laminated sand capped by wave ripples. hummocky and swaley cross-stratification are common in the southern area, near lillebælt (fig. 34). trough and tabular cross-stratified sand beds occur locally as well as tidal bundles. the cross-bedding indicates bipolar current directions towards the north-east and south-west. thin, light brown clay layers are common in the southern part. north of hvidbjerg, a dark brown, mud-dominated unit up to 3 m thick is recognised, locally capped by wood debris. the trace fossils ophiomopha isp. and skolithos isp. occur locally. the hvidbjerg member differs from the remainder of the billund formation in relation to its better sorting and its dominantly aggradational stacking pattern (e.g. rasmussen & dybkjær 2005). log characteristics. the member is characterised by low gamma-ray readings (fig. 32). high gamma-ray readings may be recorded where clay-rich, lagoonal deposits occur. fossils. a relatively rich dinocyst assemblage occurs in the hvidbjerg member (dybkjær 2004a; rasmussen & dybkjær 2005). depositional environment. deposition took place in a stormdominated shoreface environment associated with spit development, south-east of the main billund delta lobes. the core of a spit system crops out at hvidbjerg. north of hvidbjerg, shoreface sands alternate with mud-rich lagoonal deposits (fig. 21). tidal inlet deposits are observed at dykær and pjedsted where floodand ebb-dominated systems, respectively, are recorded (fig. 30). boundaries.the member overlies the vejle fjord formation; the lower boundary is marked by a change from black, organic-rich, clayey silt to white sand. at hvidbjerg, the lower boundary is erosional (fig. 31). where the member is superimposed on the skansebakke member of the vejle fjord formation, the lower boundary is identified by a change from yellowish sands of the skansebakke member to white sands of the hvidbjerg member. on the gammaray log, the boundary is characterised by a distinct shift towards low gamma-ray readings. the upper boundary is placed at the change from sandrich deposits of the hvidbjerg member to the predominantly dark brown, silty clay of the klintinghoved formation. the boundary is often erosional and overlain by a gravel lag or sand layer showing a fining-upward trend; the gravel lag commonly contains of clasts up to 4 cm in diameter. the base of the gravel lag forms the upper boundary of the member. the gamma-ray readings may be characterised by an abrupt increase followed by a gradual decrease in values or a marked decrease in gamma-ray readings sucfig. 30. tidal bundles in the hvidbjerg member exposed at pjedsted, north-west of fredericia (fig. 1). the cross-bedding dips towards the south-west and thus reflects ebb current flow. note the clay drapes and preserved bottom sets (arrows) recording sedimentation during neap tides. the cross-bedded section thus represents a neap–spring–neap cycle (i.e. c. one and a half months). the section is 1.5 m high. photograph courtesy of ole rønø clausen. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 37 38 hvidbjerg outcrop 0 1 2 3 4 5 6 7 8 9 10 m 10 11 12 13 14 15 16 17 18 19 20 cl si f mc p sand 20 21 22 23 24 25 27 26 28 cl si f mc p sand cl si f mc p sand l o w er m io ce n e b ill u n d f m v ej le f jo rd f m h vi d b je rg m b l o w er m io ce n e b ill u n d f m h vi d b je rg m b l o w er m io ce n e b ill u n d f m h vi d b je rg m b fig. 31. type section of the hvidbjerg member from the coastal exposures at hvidbjerg, south-east of vejle (fig. 1); for legend, see fig. 8, p. 17. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 38 39 ceeded by a gradual increase (e.g. andkær borehole, fig. 32); the boundary is placed at the lowest gamma-ray response. distribution. the hvidbjerg member is present in east jylland and has also been found at søndbjerg in northwest jylland (fig. 10c). biostratigraphy. the chiropteridium galea and the homo tryblium spp. dinocyst zones of dybkjær & piasecki (2010) are recorded in the hvidbjerg member. geological age. the hvidbjerg member is of aquitanian (earliest early miocene) age. fig. 33. the c. 27 m of white sand exposed at hvidbjerg represents deposition on a spit system east of the main delta lobe of the billund formation. note the stratification defined by the most bioturbated parts of the succession; photograph illustrates the upper levels of the hvidbjerg member shown in fig. 31. andkær borehole dgu no. 125.2017 90 80 70 60 50 m.b.s. gr l o w er m io ce n e b ill u n d f m k lin ti n gh o ve d f m v ej le f jo rd f m h vi d b je rg m b sk an se b ak ke m b 20 cps 100 fig. 32. reference section of the hvidbjerg member in the andkær borehole (79–58 m); for legend, see fig. 8, p. 17. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 39 addit member new member history. sand-rich fluvial and coal-bearing deposits in the silkeborg area were first studied by hartz (1909). he correlated the succession with lower miocene coal-bearing deposits in schleswig-holstein. l.b. rasmussen (1961) indicated that the fluvio-deltaic sediments of the silkeborg–skanderborg area could be of similar age to the ribe formation as defined from the arnum-1 well in southern jylland. studies of the succession in gravel pits south of silkeborg were carried out during the 1970s and 1980s, focusing on the depositional environment and diagenesis (friis 1976, 1995; hansen 1985; hansen 1995; jesse 1995). these studies referred the deposits to the middle miocene odderup formation, although friis (1995) was aware of the problems inherent in this correlation. re-investigation of the salten inland cliff and the gravel pits at addit and voervadsbro, including biostratigraphic analysis based on dinocysts, revealed that the succession is early miocene in age and should be correlated with the vejle fjord formation – billund formation depositional phase (e.s.rasmussen et al. 2006). name. after the village of addit, south-south-west of århus (fig. 1). type and reference sections. the type section of the addit member is defined as the dansand gravel pit at addit (56°02´26.33´´n, 9°37´59.62´´e; fig. 35). the member is also exposed in the voervadsbro gravel pit which forms the primary reference section (fig. 36), and in the inland cliff at salten. the secondary reference section is the interval from 117 to 55 m (119–55 m md) in the addit mark borehole (dgu no. 97.928; fig. 37). thickness. in the type section, the addit member is over 33 m thick; neither top nor base are seen. in the addit borehole nearby (fig. 1), the member is 50 m thick (plate 6) and 62 m was penetrated in the borehole at addit mark (fig. 37; plate 6). where well developed in central and north-east jylland, the member is typically 20–50 m thick (see plates 1, 2, 5, 6). lithology. the succession is typically composed of two sandand gravel-rich units separated by fine-grained, sandy and clayey sediments, commonly with intercalated coal layers (figs 35, 36, 38; plates 1, 6), though the middle heterogeneous unit may be absent or poorly developed. the sands consist almost solely of quartz and quartzitic sandstone lithic grains, with minor content of mica and heavy minerals; clasts of well-rounded chert may occur. the two sandrich units are characterised by fining-upward trends and possess sheet geometry. the lower part of each unit consists of trough cross-stratified, coarse-grained sand and gravel alternating with large-scale cross-stratified sand (figs 35, 36, 39). upwards, these sand-rich units are progressively dominated by tabular co-sets of cross-stratified sand. the sand-rich succession may be capped by fine-grained, crossbedded sand showing lateral accretion structures. the coalbearing, fine-grained sand and clay layer, sandwiched between the coarser units, consists of cross-bedded sand and alternating thin, rippled, fine-grained sand and clay layers. bioturbation is observed rarely. wood fragments are abun40 fig. 34. hummocky cross-stratified sand of the billund formation (hvidbjerg member) overlying interbedded, hummocky cross-stratified sands and clays of the vejle fjord formation; hindsgavl, near middlefart (fig. 1). bulletin 22_ gsb191-indhold 04/03/11 12.41 side 40 41 addit outcrop 0 1 2 3 4 5 6 7 8 9 10 m 10 11 12 13 14 15 16 17 18 19 20 cl si f mc p sand 20 21 22 23 24 25 26 27 28 29 30 30 31 32 33 cl si f mc p sand cl si f mc p sand cl si f mc p sand b ill u n d f m l o w er m io ce n e a d d it m b b ill u n d f m l o w er m io ce n e a d d it m b b ill u n d f m l o w er m io ce n e a d d it m b b ill u n d f m l o w er m io ce n e a d d it m b fig. 35. type section of the addit member in the addit gravel pit, south-east of silkeborg (fig. 1); for legend, see fig. 8, p. 17. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 41 42 voervadsbro outcrop 0 1 2 3 4 5 6 7 8 9 10 m 10 11 12 13 14 15 16 17 18 19 20 cl si f mc p sand 20 21 22 23 26 25 24 cl si f mc p sand cl si f mc p sand b ill u n d f m a d d it m b b ill u n d f m a d d it m b b ill u n d f m l o w er m io ce n e l o w er m io ce n e l o w er m io ce n e a d d it m b fig. 36. primary reference section of the addit member in the voervadsbro gravel pit, south-east of silkeborg (fig. 1), where the lower part of the member is exposed; for legend, see fig. 8, p. 17. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 42 dant at certain horizons and petrified wood is common at voervadsbro (weibel 1996). log characteristics. the member is characterised by low gamma-ray readings, especially in the lower part (fig. 37). the sand-rich part is commonly characterised by a slight upward increase in gamma-ray values. a moderate–high gamma-ray response commonly characterises the middle part of the member, reflecting the clay-rich and coal-bearing deposits at this level. fossils. the addit member contains fossil wood (weibel 1996), leaves and seeds (ravn 1907). dinocysts occur very sporadically in the addit member (dybkjær 2004a, b; e.s. rasmussen et al. 2006). depositional environment. the lower sands of the member were deposited as migrating three-dimensional dunes (main channel) and migrating unit and compound bars in a braided fluvial system (hansen 1985; hansen 1995; e.s. rasmussen et al. 2006). the upper part of the member was deposited as migrating two-dimensional dunes; sedimentary structures such as cross-bedded sand beds with preserved bottomsets and normally graded foresets indicate tidal influence (pontén & plink-björklund 2007). the upper part of the sand succession, showing lateral accretion, was laid down in a point bar of a meandering fluvial system. the fine-grained middle part of the member was deposited in a flood plain and lake environment that was occasionally flooded by the sea as indicated by the rare presence of dinocysts and ophiomorpha trace fossils. boundaries. where the addit member directly overlies the vejle fjord formation, the lower boundary is marked by an abrupt change from dark brown, silty clay or clayey silt to grey, coarse-grained sand and gravel (fig. 37; plates 1, 2, 6). where the addit member overlies the hvidbjerg member, the boundary is marked by an erosional boundary where white, fineto medium-grained sand is overlain by gravel (fig. 29). on the gamma-ray log, the lower boundary is shown as a prominent shift on the gamma-ray log where the addit member overlies the vejle fjord formation. where the member overlies the hvidbjerg member, the lower boundary is placed at the change of gamma-ray reading from a gradual upward decrease in gamma-ray readings to consistently low or decreasing-upward gamma-ray readings. the upper boundary is placed at the change from sandrich deposits of the addit member to predominantly dark brown, silty clays of the klintinghoved formation. the boundary may be erosional, being overlain by a gravel lag or sand layer showing a fining-upward trend; the gravel lag commonly contains clasts up to 4 cm in diameter. the base of the gravel lag or sand layer forms the upper boundary of the member. the gamma-ray log is commonly characterised by an abrupt increase in gamma-ray values (e.g. 43 fig. 37. secondary reference section of the addit member in the addit mark borehole (117–55 m); for legend, see fig. 8, p. 17. k.: klintinghoved. 140 0 cps 90 130 120 110 addit mark borehole dgu no. 97.928 100 90 80 70 60 m.b.s. gr v ej le f jo rd f m l o w er m io ce n e b ill u n d f m k . f m a d d it m b bulletin 22_ gsb191-indhold 04/03/11 12.41 side 43 resen, plate 5; isenvad, plate 6). locally, a decrease in gamma-ray readings is succeeded by a gradual increase in gamma-ray values (e.g. hammerum, sunds, plate 2). here the boundary is placed at the lowest gamma-ray readings. distribution. the addit member is found in the central and northern parts of jylland. it is especially well developed in the area south of silkeborg, in an elongate zone striking from resen (south-west of viborg) to the area between herning and ikast (figs 1, 10c). biostratigraphy. the homotryblium spp. dinocyst zone of dybkjær & piasecki (2010) is recorded in the addit member. geological age. the addit member is of aquitanian (earliest early miocene) age. 44 fig38 coal bedcoal bed quaternaryquaternary fig. 38. addit member at the addit gravel pit showing the two finingupward sandand gravel-rich units and the intercalated coal unit. the height of the section is 40 m. fig. 39. cross-bedded sand and gravel of the addit member in the addit gravel pit deposited as a mid-channel bar in a braided fluvial system. the height of the section is 3 m. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 44 klintinghoved formation redefined formation general. the marine clay-rich deposits of the klintinghoved formation interdigitate towards the north-east with the more proximal sand-rich deltaic sediments of the bastrup formation. these formations thus alternate up-section in certain boreholes (e.g. plates 1, 2, 6). history. the mollusc fauna of marine clay-rich deposits at klintinghoved was described by sorgenfrei (1940). the deposits were defined as the klintinghoved formation in a later publication (sorgenfrei 1958). the klintinghoved formation was included in the stratigraphy of l.b. rasmussen (1961). name. after klintinghoved cliff, flensborg fjord (figs 1, 42). type and reference sections. following sorgenfrei (1958), the type section is the outcrop at klintinghoved cliff (54°53´23.15´´n, 9°49´43.62´´e; figs 40, 42). the reference section is designated in the cored sdr. vium borehole (dgu no. 102.948; 54°53´23.18´´n, 9°49´43.94´´e) from 288 to 132 m (figs 40, 41). thickness. at klintinghoved, the exposed section is 3.5 m thick; neither base nor top of the formation is seen. in the subsurface, the formation is 10–50 m thick in central jylland, thickening to over 125 m in the west and southwest (e.g. sdr. vium, fig. 40). lithology. the formation consists of dark brown, silty clay with subordinate intercalated sand beds (figs 43, 44). the sand beds are sharp based and homogenous to finely laminated; double clay layers are recognised locally. in the cored borehole at sdr. vium, the formation is dominated by dark brown mud with intercalated sand beds (figs 40, 41, 45). the sand beds typically show sharp lower boundaries, and are commonly structureless in the lower part passing upward into laminated sand. log characteristics. the formation is characterised by moderate to high gamma-ray values (fig. 40). the log pattern is highly serrated, reflecting interbedded muds and sands at various levels, and shows a general decrease in gammaray response upwards. fossils. the klintinghoved formation contains a rich mollusc fauna (sorgenfrei 1958). shark teeth also occur and marine microfossils such as foraminifers (laursen & kristoffersen 1999) and dinocysts (dybkjær & rasmussen 2000; dybkjær 2004a; rasmussen & dybkjær 2005) are abundant and diverse. 45 bulletin 22_ gsb191-indhold 04/03/11 12.41 side 45 46 klintinghoved outcrop 0 1 2 3 m cl si f mc p sand sdr. vium borehole dgu no. 102.948 300 200 m.b.s. 200 190290 280 270 260 250 240 230 220 210 cl si f mc p sand not cored no gr log available cl si f mc p sand l o w er m io ce n e k lin ti n gh o ve d f m l o w er m io ce n e k lin ti n gh o ve d f m l o w er m io ce n e k lin ti n gh o ve d f m 160 150 140 130 120 110 l o w er m io ce n e a rn u m f m b as tr u p f m k lin ti n gh o ve d f m gr 40 100 cps fig. 40. type and reference sections of the klintinghoved formation. the type section is defined at klintinghoved, east of sønderborg, where 3.5 m of the formation is exposed. the reference section is the interval from 288 to 132 m in the sdr. vium borehole. the intervals outlined in red are shown in detail in fig. 41; for legend, see fig. 8, p. 17. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 46 47 depositional environment.the klintinghoved formation was deposited in shelf, delta slope and lower shoreface environments. water depths were in the order of 15 to 60 m, but locally up to 100 m based on the height of clinoforms seen on seismic data. the depositional environment was strongly influenced by storms and tidal processes. boundaries. the lower boundary is, in the northern part, characterised by a change from sand-rich deposits of the billund formation to the predominantly dark brown, silty clay of the klintinghoved formation (e.g. store vorslunde, fasterholt, sunds and resen boreholes, plate 2). in the southern part where the klintinghoved formation overlies the vejle fjord formation, the boundary is not marked by significant changes in lithology, although the vejle fjord formation tends to be slightly more consolidated. on the gamma-ray log, a weak to marked shift to higher gammaray values defines the boundary, especially where the klintinghoved formation overlies the billund formation (e.g. rødding, almstok, store vorslunde, fasterholt, sunds and resen boreholes, plate 2). at outcrop, the lower boundary is often erosional and overlain by a gravel lag, as seen at rønshoved and børup (rasmussen & dybkjær 2005); the gravel lag commonly contains clasts up to 4 cm in di ameter. the upper boundary is either sharp, exemplified by the bastrup borehole (fig. 52; plate 8), or gradational as in the almstok borehole (fig. 52; plate 2). in the bastrup borehole, the upper boundary is placed where grey mud is sharply overlain by grey, medium-grained sand. in boreholes where a more gradational development occurs, the boundary is marked by a change from alternating beds of sand and mud to a clean sand unit at least 5 m thick and comprising at least 75% sand. on the gamma-ray log, the 275 274 273 272 271 270 269 268 267 266 265 cl si f mc p sand 140 139 138 137 136 135 134 133 cl si f mc p sand sdr. vium borehole dgu no. 102.948 sdr. vium borehole dgu no. 102.948 m.b.s m.b.s fig. 41. detailed sedimentological logs of representative intervals of the klintinghoved formation in the reference section (for location, see fig. 40; for legend, see fig. 8 on page 17). bulletin 22_ gsb191-indhold 04/03/11 12.41 side 47 48 double clay layersdouble clay layers fig. 42. the klintinghoved cliff, viewed from the east (cliff is c. 10 m high). the location of the type section of the klintinghoved formation is arrowed. fig. 43. alternating clay and bioturbated and laminated sand of the klintinghoved formation at the type locality. note the double clay layers in the sand indicating tidal influence on deposition. fig. 44. interlaminated, dark brown clayey silt and thin, fine-grained sand of the klintinghoved formation at the type locality. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 48 boundary is generally characterised by a minor decrease in gamma-ray values followed by a consistent decrease in values upwards, as seen in the almstok borehole (fig. 52; plate 2) and in the holstebro and klosterhede boreholes (plate 4). in the bastrup borehole, the upper boundary is characterised by a distinct decrease in gamma-ray values. in western jylland where klintinghoved formation is overlain by the arnum formation (e.g. kvong, sdr. vium boreholes, plate 4), the boundary is placed at a distinct increase in gamma-ray readings separating coarsening-upward units of the klintinghoved and arnum formations. distribution. the klintinghoved formation is distributed in the northern part of central jylland and in western and southern jylland (fig. 10d). biostratigraphy. the thalassiphora pelagica and sumatradinium hamulatum dinocyst zones of dybkjær & piasecki (2010) are recognised in the klintinghoved formation. geological age. the klintinghoved formation is of late aquitanian to early burdigalian (early miocene) age. subdivision. the klintinghoved formation includes the new kolding fjord member. 49 fig. 45. core sections from the sdr. vium borehole, illustrating interbedded, dark brown silty clays and sharp-based sands of the klintinghoved formation. a: 272.70 m (base of illustrated section); b: 250.45 m (base). the sand beds are normally graded and homogenous to weakly laminated in the lower part. note the double clay layers (b, arrow) indicating tidal influence on sedimentation. 1 cm a b bulletin 22_ gsb191-indhold 04/03/11 12.41 side 49 kolding fjord member new member history. sand and organic-rich clayey sediments exposed at lillebælt were studied by radwanski et al. (1975), e.s. rasmussen (1995) and friis et al. (1998). in these studies, the sediments were referred tentatively to the vejle fjord formation of previous usage. however, a biostratigraphic study by dybkjær & rasmussen (2000) revealed that the sediments were significantly younger than the vejle fjord formation (as recognised here) and equivalent in age to the klintinghoved formation (l.b. rasmussen 1961). name. the kolding fjord member crops out at a number of localities along lillebælt and kolding fjord. it is named after kolding fjord, where the type locality of rønshoved is situated. type and reference sections. the type section is the exposure at rønshoved on the southern side of kolding fjord (55°29´26.90´´n, 9°38´36.10´´e; figs 1, 46). other localities where the member is exposed are hagenør, børup, galsklint and fænø in the lillebælt and kolding fjord area (fig. 1). a minor outcrop is also recognised at gyldendal, limfjorden. the reference section is the outcrop at hagenør (figs 1, 47). thickness. the kolding fjord member is 11 m thick at rønshoved and c. 8 m at hagenør (figs 46, 47). although rarely exceeding 10 m, developments up to 20 m thick are recognised locally (e.g. vonsild and vind boreholes, plates 1, 8). lithology. the member is composed of white to yellow, fineto medium-grained sand with a few thin, brown clay layers. at the type section, the basal unit is a gravel layer c. 10 cm thick that contains clasts up to 4 cm in diameter. the clasts consist of almost pure quartz and quartzitic sandstone. the succeeding sandy part of the member in the type section is dominated by hummocky and swaley crossstratified silt and fine-grained sand (figs 46, 48). the more clayey part is dominated by heterolithic mud which shows hummocky cross-stratification and clear rhythmicity i.e. double clay layers and alternating sandand mud-rich units. layers up to 2 m thick of dark brown, organic-rich, clayey silt may be intercalated in the sand (figs 49, 50). homo50 rønshoved (east) outcrop 0 1 2 3 4 5 6 7 8 9 10 m cl si f mc p sand k lin ti n gh o ve d f m l o w er m io ce n e k o ld in g fj o rd m b b ill u n d f m h vi d b je rg m b 10 11 12 13 cl si f mc p sand k lin ti n gh o ve d f m l o w er m io ce n e k o ld in g fj o rd m b fig. 46. type section of the kolding fjord member at rønshoved, east of kolding; for legend, see fig. 8, p. 17. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 50 geneous sand beds capped by wave-ripples are also common on top of lagoonal deposits (fig. 51); wave-ripple crests are oriented north-west–south-east. trace fossils, including macaronichnus isp., ophiomorpha nodosa and echinoid burrows, are common in the kolding fjord member (radwanski et al. 1975). log characteristics. the member is characterised by low to moderate gamma-ray values. the log pattern is serrated; high gamma-ray values are registered where lagoonal, clay-rich deposits dominate. fossils. the kolding fjord member contains a dinocyst assemblage of variable richness (dybkjær & rasmussen 2000; rasmussen & dybkjær 2005). depositional environment. deposition took place on a stormdominated coast in a lower and upper shoreface environment (friis et al. 1998; rasmussen & dybkjær 2005). the fine-grained, heterolithic part was deposited in a lagoonal environment with some tidal influence. the upper part of the member was deposited as washover fans on the back-barrier flat during the final degradation of the barrier complex. boundaries. in the type section, the lower boundary is erosional and, as in other exposures (e.g. børup, galsklint) and borehole sections (e.g. stakroge, plate 3), is defined by a distinct change from the sandy deposits of the billund formation to gravel-dominated layers of the lowermost kolding fjord member. in such cases, the gamma-ray log shows a marked decrease in values at the boundary (plate 3). in the vonsild borehole, however, located near the type and reference sections, the lower boundary is recognised by a prominent increase in gamma-ray readings, due to the presence of fine-grained, lagoonal sediments in the lower part of the member (plates 1, 8). it is acknowledged that identification of this boundary may be difficult on the gamma-ray log where shoreface sands occur both beneath and above the boundary and the transgressive lag is thin. the upper boundary is characterised by a change from the sand-dominated succession of the kolding fjord member to dark brown, clayey silts of the klintinghoved formation. the gamma-ray log shows a distinct increase in gamma-ray values. distribution. the member is recognised in east jylland and south-west of holstebro in west jylland (fig. 10d). biostratigraphy. the thalassiphora pelagica and sumatradinium hamulatum dinocyst zones of dybkjær & piasecki (2010) occur in the kolding fjord member. geological age. the kolding fjord member is of late aquitanian to early burdigalian (early miocene) age. 51 hagenør outcrop 0 1 2 3 4 5 6 7 8 m cl si f mc p sand k lin ti n gh o ve d f m l o w er m io ce n e q u at . k o ld in g fj o rd m b fig. 47. reference section of the kolding fjord member at hagenør; for legend, see fig. 8, p. 17. quat.: quaternary. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 51 52 fig. 48. heterolithic deposits of the kolding fjord member sharply overlain (at 5.3 m in fig. 46) by hummocky cross-stratified sand at rønshoved in the type section. the heterolithic succession is characterised by alternating hummocky cross-stratified sand and sandy clay and various types of ripple-laminated sand. about 3 m of the section is shown. fig. 49. exposure of the kolding fjord member in the reference section at hagenør. the lower part of the hagenør outcrop is characterised by two organic-rich, clayey silt deposits separated by bioturbated sand (see fig. 47). the upper part of the exposure is dominated by alternating sand and clay layers; the sand beds are typically sharp based, homogenous to weakly laminated in the lower part and capped by waveor current-ripples. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 52 bastrup formation new formation general. the new bastrup formation is recognised primarily in the subsurface. this fluvio-deltaic, sand-dominated formation interdigitates in a complex manner with the more distal, marine, mud-rich klintinghoved formation. these formations thus alternate up-section in certain boreholes (e.g. plates 1, 2, 6). name. after bastrup village, south-west of kolding (fig. 1). type and reference sections. the type section is the interval from 108 to 84 m (110–84 m md) in the bastrup borehole (dgu no. 133.1298; 55°24´21.58´´n, 9°14´47.40´´e; fig. 52). the reference section is the interval from 160 to 111 m (160–111 m md) in the borehole at almstok (dgu no. 114.1858; fig. 52). thickness. the thickness of the bastrup formation is 24 m in the type section, but the formation is commonly c. 50 m thick (see reference section, fig. 52 and plates 2, 3). a maximum thickness of 100 m was penetrated in the borehole at løgumkloster (plate 3). lithology. the bastrup formation consists predominantly of grey, mediumto coarse-grained sand with intercalated gravel layers; the diameter of gravel clasts rarely exceeds 2 cm. petrologically, the sand is dominated by quartz and quartzite lithic grains with minor content of mica and heavy minerals. in a few boreholes, however, a high concentration of mica has been recorded (e.g. estrup). dark brown, organic-rich, silty clay is locally present. the formation is characterised by both coarsening-upward and fining-upward depositional patterns. the upper part of the formation is commonly characterised by a 15–30 m thick fining-upward succession consisting of coarse-grained to finegrained sand. in the north, gravel commonly forms the 53 fig. 50. close-up of lagoonal facies in the kolding fjord member at hagenør. the light brown deposits that are capped by sand ripples and sandwiched between dark lagoonal clays contain marine palynomorphs and represent a short marine incursion; the strike of the ripple crests is nw–se. the illustrated section is c. 2 m high. fig. 51. close-up of the alternating sand and clay layers of the kolding fjord member exposed in the upper part of the hagenør reference section. spade handle for scale. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 53 base of the fining-upward units. clay-rich sediments with subordinate intercalations of coal are often sandwiched between sand-rich units. log characteristics. the formation is characterised by low gamma-ray values (fig. 52). the log pattern is serrated and shows both decreasingand increasing-upward trends through the succession. the decreasing trend is associated with delta progradation and the increasing-upward trend is associated with channel-fill deposits (i.e. point bars) which are common in the upper levels of the formation, and can locally be demonstrated on seismic data (e.s. rasmussen et al. 2007). fossils. a sparse foraminifer assemblage occurs in the distal part of the bastrup formation (laursen & kristoffersen 1999). the dinocyst flora is variable overall, being rich at some levels and very sparse/impoverished at other levels (dybkjær 2004a; dybkjær & piasecki 2010). depositional environment. deposition took place in deltaic and fluvial environments. well developed point bars and fluvial channels are common in the upper part (e.s. rasmussen et al. 2007; e.s. rasmussen 2009b). the intercalated mud represents floodplain deposition. boundaries. the lower boundary is either sharp, for example in the type section of the bastrup borehole or gradational as in the almstok reference section (fig. 52). in the type section, the lower boundary is placed where grey mud is sharply overlain by grey, medium-grained sand; on the gamma-ray log, this lower boundary is defined at a marked decrease in gamma-ray values. a gravel layer is commonly present at the base of the bastrup formation. in gradational sections showing interbedded sands and muds, becoming sandier upwards, the boundary is defined at the base of the first significant sand interval (at least 5 m thick) in which the sand to mud ratio is greater than 75%. in such gradational sections, the log response reflects the transitional nature of the boundary, showing a minor decrease in gamma-ray values followed by a consistent overall decrease upwards (e.g. the almstok borehole, fig. 52). the upper boundary is defined by a sharp transition from grey and white sand of the bastrup formation to dark brown, silty clay of the arnum formation. in central jylland, the arnum formation is developed as a grey to white silt, which rests with a sharp boundary on the sand-rich bastrup formation. on the gamma-ray log, this upper boundary is typically identified by a prominent shift to higher values. distribution. the formation is present in southern and central jylland. towards the north-east, the formation is truncated and it pinches out towards the south-west (fig. 10e). biostratigraphy. the sumatradinium hamulatum and cordosphaeridium cantharellus dinocyst zones of dybkjær & piasecki (2010) occur in the bastrup formation. geological age. the bastrup formation is of early burdigalian (early miocene) age. subdivision. the bastrup formation includes the new resen member. 54 bastrup borehole dgu no. 133.1298 120 110 100 90 80 70 m.b.s. gr gr almstok borehole dgu no. 114.1858 170 160 150 140 130 120 110 100 m.b.s. l o w er m io ce n e b as tr u p f m r es en m b k lin ti n gh o ve d f m a r. fm o d d er u p. f m st au n in g m b l o w er m io ce n e b as tr u p f m k lin ti n gh o ve d f m a rn u m f m r es en m b 0 cps 120 0 150 cps fig. 52. type and reference sections of the bastrup formation. the type section is from 108 to 84 m in the bastrup borehole. the reference sec tion is from 160 to 111 m in the almstok borehole; for legend, see fig. 8 on page 17. ar.: arnum. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 54 resen member new member general. this member is recognised widely in the bastrup formation, representing fluvial-dominated facies that commonly are inferred to be incised into the undifferentiated bastrup formation deltaic facies. it is mainly recognised in the subsurface, but coal was formerly mined in a pit near resen, south of skive (fig. 1). name. after the village of resen, south of skive, where a brown-coal pit was mined (fig. 1). type and reference sections. the type section is the composite interval from 124 to 112 m and from 97 to 70 m (125–113 m md, 97–71 m md) in the borehole at hammerum, east of herning (dgu no. 85. 2429; 56°07´55.45´´n, 9°05´33.52´´e; figs 1, 53). the reference section is the interval from 104 to 67 m (105– 67 m md) in the egtved borehole, south-west of vejle (dgu no. 124.1159; figs 1, 53). thickness. the member is 39 m thick in the type section (fig. 53) and is typically in the range 10–40 m thick (plates 2, 3, 6, 7). lithology. the member consists of grey, mediumto coarsegrained sand with intercalated gravel layers. dark brown, organic-rich, silty clay with some coal is present locally. the member is typically characterised by 10–30 m thick fining-upward successions; a number of boreholes show stacked, fining-upward cycles that may be separated by intervals referred to the bastrup formation (undifferentiated). log characteristics. the member is characterised by low gamma-ray readings. the log pattern is serrated and, where simply developed (e.g. billund, plate 2), shows an increasing trend upwards, reflecting the origin of these sand-rich units as channel fill deposits. in some boreholes, such channel sands are separated by finer-grained deposits showing moderate–high gamma-ray values (fig. 53). fossils. the dinocyst flora is variable overall, being rich at some levels and very sparse/impoverished at other levels (dybkjær 2004a; dybkjær & piasecki 2010). depositional environment. deposition took place in fluvial environments, and well-developed point bars and fluvial channels are common (e.s. rasmussen et al. 2007; e.s. rasmussen 2009b). the intercalated mud represents floodplain deposition and some marine influence has also been recognised, especially in the southern part. the most extensive coal formation was within freshwater lakes and mires developed in the rim synclines around salt diapirs, e.g. the sevel and mønsted salt structures south of skive (japsen & langtofte 1991). boundaries. the resen member is bounded both by sandrich units (e.g. bastrup formation beneath, vandel member above) and by mud-rich units (klintinghoved formation beneath, arnum formation above). where succeeding the klintinghoved formation, the boundary is sharp and placed where dark brown, clayey silts of the klintinghoved formation are sharply overlain by grey, mediumto coarsegrained sands, locally with a basal gravel layer. on the gamma-ray log, this relationship is recorded by an abrupt shift to lower values (fig. 53). in sections where the resen member succeeds the undifferentiated bastrup formation, the boundary is defined at the base of coarser sand/gravel deposits at a shift from decreasing-upward gamma-ray values (bastrup formation) to increasing-upward gamma-ray values (resen member channel sands). 55 fig. 53. type and reference sections of the resen member. the composite section (124–112 m, 97–70 m) in the hammerum borehole is designated as the type section. the reference section is the interval from 104 to 67 m in the egtved borehole; for legend, see fig. 8, p. 17. a.: arnum. klint.: klintinghoved. od.: odderup. res.: resen. hammerum borehole dgu no. 85.2429 130 cps 120 110 100 90 80 70 60 m.b.s. egtved borehole dgu no. 124.1159 110 20 cps 80 100 90 80 70 60 m.b.s. grgr b as tr u p f m o d . f m l o w e r m io ce n e r e se n m b r e s. m b b as tr u p f m a . f m k lin t. f m k lin t. f m l o w e r m io ce n e r e se n m b 0 60 bulletin 22_ gsb191-indhold 04/03/11 12.41 side 55 56 the upper boundary is typically defined by a sharp transition from grey and white sands of the resen member to dark brown, silty clay of the arnum formation; on the gamma-ray log this is reflected by an abrupt increase in values. in central jylland, the upper boundary is characterised by a sharp change from grey and white sand to grey and white silt of the vandel member. on the gamma-ray log, this facies transition is reflected by an increase in gamma-ray values that continues up through the vandel member. distribution. the member is present in southern and central jylland (fig. 10e; plates 1, 2, 6). towards the northeast, the member is truncated and it pinches out towards the south-west. biostratigraphy. the sumatradinium hamulatum and cordosphaeridium cantharellus dinocyst zones of dybkjær & piasecki (2010) occur in the resen member. geological age. the resen member is of early burdigalian (early miocene) age. arnum formation revised formation general. the marine clay-dominated arnum formation is only recognised in the subsurface where it shows complex interdigitation with the nearshore sand-rich odderup formation. these two formations thus altenate up-section in some boreholes (plates 1–9). history. the arnum formation was defined by sorgenfrei (1958) to encompass the dark micaceous marine clays occurring stratigraphically above the ribe formation (of previous usage). name. after arnum village in southern jylland (fig. 1). type and reference sections. the arnum formation was penetrated in two boreholes at arnum (dgu no.150.13, dgu no.150.25b; both at 55°14´48.07´´n, 8°58´18.48´´e) from 107 to 40 m and from 107.5 to 40 m respectively (sorgenfrei 1958); together these sections form the type section. the composite interval 132–111 m and 98–51 m in the cored borehole, sdr. vium (dgu no. 102.948) is designated as the reference section (fig. 54). a secondary reference section is defined as the interval from 55 to 37 m (56–39 m md) in the store vorslunde borehole (dgu no. 104.2325; fig. 55). thickness. the formation is c. 93 m thick in the type borehole (sorgenfrei 1958). the formation is commonly only a few tens of metres thick in the north-east of the area but thickens west and south (plates 4, 7, 9); about 130–150 m were encountered in the borg-1 and rømø boreholes and nearly 200 m in the forumlund borehole (plate 4). lithology. the arnum formation consists of dark brown, silty clay with occasional shell beds. thin laminated, finegrained sand beds are common. the sand beds commonly display a sharp lower boundary succeeded by laminated and low-angle cross-bedded sand capped by wave laminated sand. micro-hummocky cross-stratification is common. some of the wave-rippled sand beds have sharp erosive upper boundaries overlain by mud (fig. 56a–c). thin sand beds and silt layers may have a high content of heavy minerals; glaucony is present and locally forms discrete lamina (fig. 56d). log characteristics. the formation is characterised by moderate–high gamma-ray values (figs 54, 55). the log pattern is serrated (reflecting subordinate interbedded sands) and commonly shows an overall decreasing trend upwards. discrete gamma-ray peaks may be related to silt and sand beds rich in heavy minerals. fossils. the arnum formation contains a rich assemblage of marine molluscs (sorgenfrei 1958; l.b. rasmussen 1961). rich foraminifer and dinocyst assemblages also occur in this formation (laursen & kristoffersen 1999; dybkjær & piasecki 2010). depositional environment. the arnum formation was deposited in a fully marine shelf environment. the water depth is unknown, but the concentration of heavy minerals and the presence of scours and wave-rippled sand may indicate rather shallow water with frequent reworking and sorting of sediments. boundaries. the arnum formation is typically bounded by sand-rich formations, the bastrup formation or vandel member beneath and the odderup formation, both beneath and above. in the former case, the lower boundary is defined by a sharp transition from grey and white sand of the bastrup formation to dark brown, silty clay of the arnum formation, recorded on the gamma-ray log as an abrupt increase in values. in central jylland, the lower boundary is defined at the change from grey and white sand to grey bulletin 22_ gsb191-indhold 04/03/11 12.41 side 56 and white silt of the vandel member. in western and southern sections, the bastrup formation is absent and the clayrich arnum formation succeeds silty clays of the klintinghoved formation (fig. 54). this boundary may be difficult to locate but is typically placed where the consistently decreasing-upward gamma-ray trend of the uppermost klintinghoved formation is succeeded by the ‘noisy’, serrated pattern of the arnum formation (e.g. fig. 54; kvong borehole, plate 4). the upper boundary is placed at the base of the first significant occurrence of grey fine-grained sand, commonly with a high content of heavy minerals, that is thicker than 5 m with a sand/mud ratio of at least 75%. on the gammaray log, the upper boundary with the stauning member may be difficult to recognise but is marked by a shift from serrated and moderate–high gamma-ray values to low–moderate values, albeit still serrated in nature (e.g. stauning borehole, plate 6). at rømø, in the far south-west (plate 9), the odderup formation is absent and the arnum formation is overlain by the hodde formation (måde group); the boundary is placed at the shift from consistent moderate–high gamma-ray values to increasing-upward values. 57 100 90 80 70 60 50 60 m.b.s. m 160 150 140 130 120 110 cl si f mc p sand cl si f mc p sand l o w er m io ce n e l o w er m io ce n e a rn u m f m a rn u m f m a rn u m f m o d d er u p f m st au n in g m b b as tr u p f m k lin ti n gh o ve d f m sdr. vium borehole dgu no. 102.948 121 120 m.b.s. 127 126 125 124 123 122 cl si f mc p sand sdr. vium borehole dgu no. 102.948 40 cps 100 store vorslunde borehole dgu no. 104.2325 60 50 40 30 m.b.s. gr l o w er m io ce n e a rn u m f m o d d er u p f m b as . f m 0 cps 80 fig. 55. the secondary reference section for the arnum formation is the interval from 55 to 37 m in the store vorslunde borehole; for legend, see fig. 8, p. 17. bas.: bastrup. fig. 54. the primary reference section of the arnum formation is the composite interval (132–111 m, 98–51 m) in the cored sdr. vium borehole; for legend, see fig. 8, p. 17. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 57 distribution. the formation is recognised in jylland, southwest of a line from struer to horsens (figs 1, 10f). biostratigraphy. the cordosphaeridium cantharellus, exochosphaeridium insigne, cousteaudinium aubryae and labyrinthodinium truncatum dinocyst zones of dybkjær & piasecki (2010) occur in the arnum formation. geological age. the arnum formation is of burdigalian to early langhian (early – early middle miocene) age. subdivision. the arnum formation includes the new vandel member. vandel member new member general. the lack of exposure of this member precludes detailed description and environmental interpretation. it is defined as a discrete member of the arnum formation since it forms a recognisable marker interval between the coarse siliciclastics of the bastrup formation beneath and the mud-rich facies of the arnum formation above. name. after the village of vandel, east of billund (fig. 1). type and reference sections. the type section is the interval from 114 to 102 m (112–102 m md) in the vandel mark borehole (dgu no.115.1371; 55°42´47.99´´n, 9°10´54.82´´e; figs 1, 57). the reference section is the interval from 100 to 97 m (100–97 m md) in the grindsted borehole (dgu no. 114.2038; fig. 57). thickness. the thickness of the member rarely exceeds the 12 m recorded in the borehole at vandel mark (plate 7). lithology. in both the vandel and the grinsted boreholes, log and/or cuttings data indicate a lowermost sand or gravel layer, fining upwards into mud-rich deposits. the diagnostic feature of the vandel member, however, is the occurrence of grey to white silt with a high content of heavy mine rals; clasts of reworked reddish eocene clay may be present. 58 fig. 56. slabbed core sections from the sdr. vium borehole showing typical lithologies of the arnum formation. a: dark brown clayey silt interbedded with hummocky cross-stratified sand; 59.20 m (base of illustrated core). b: hummocky cross-stratified sand bed bounded by dark brown silty clays; 73.40 m (base). c: heterolithic deposits showing double clay layers (arrows); 95.60 m (base). note the small-scale faults cutting the succession, possibly due to contemporaneous seismic activity. d: bioturbated clay with a 1 mm lamina rich in glaucony; 125.00 m (base). 1 cm 1 cm da glauconyglaucony b c 1 cm 1 cm bulletin 22_ gsb191-indhold 04/03/11 12.41 side 58 log characteristics. the member shows intermediate gamma-ray readings overall with subordinate low values near the base (sandy beds) and localised high peaks (? heavy mineral sands). fossils. no fossils have been recorded. depositional environment.the depositional setting is unclear but the member caps fluvio-deltaic deposits (resen member) of the bastrup formation. the absence of fossils could point towards a floodplain depositional environment. boundaries. the lower boundary is defined by a lithological shift from grey sand to grey and white silt as observed in borehole cuttings samples. this boundary is difficult to position on the gamma-ray log alone; a minor increase in values is observed in the type section (fig. 57), followed by a weak increasing-upward trend. the upper boundary is placed at the top of the interval of white to grey silt. a slight, but distinct decrease in gamma-ray values is recognised at the upper boundary in the type section. distribution. the member is recognised in central jylland (fig. 10f). biostratigraphy. the vandel member is barren of dinocysts, but the cordosphaeridium cantharellus dinocyst zone (dybkjær & piasecki 2010) occurs in the lithostratigraphic units below and above. geological age. the vandel member is of burdigalian (late early miocene) age. odderup formation redefined formation history. the odderup formation was defined by l.b. rasmussen (1961), from the borehole at odderup brickworks where the succession of brown coal and quartz sand from 40.3 to 28.2 m was defined as the type section. koch (1989) subsequently erected the fasterholt member and included this in the odderup formation. the formation is redefined here, based on the more extensive subsurface database now available, to include the marine sand-dominated succession, commonly rich in heavy minerals, that is associated with the largely terrestrial sediments recognised in the early work. name. after odderup village in western jylland (fig. 1). type and reference sections. following l.b. rasmussen (1961), the type section is the borehole at odderup (dgu no. 103.150; 55°52´19.05´´n, 8°37´42.28´´e) from 40.3 to 28.2 m. the formation is exposed at the abildaa brown coal museum near ørnhøj but only the brown-coal-bearing fasterholt member is present here. the primary reference section is the interval from 37 to 1 m (39–1 m md) in the borehole at store vorslunde (dgu no. 104.2325). the secondary reference section illustrates the alternation of the odderup and arnum formations that is observed in a number of boreholes (plates 2–9); the odderup formation is represented in the intervals from 118 to 110 m (118–111 m md) and 90 to 41 m (90– 42 m md) in the rødding borehole (dgu no. 141.1141; fig. 58). thickness. the formation is c. 12 m thick at the type section and about 36 m thick in the primary reference section. in central jylland, it commonly exceeds 40 m (plates 2, 7) and an exceptionally thick development was recorded in the tinglev borehole (c. 165 m; plates 1, 9). lithology. the formation consists of fineto coarse-grained sand with some intercalation of clay beds and brown coal. the formation consists of quartz and clast of quartzites 59 vandel mark borehole dgu no. 115.1371 grindsted borehole dgu no. 114.2038 120 110 100 90 m.b.s. gr gr 110 0 100 100 90 80 m.b.s. b as tr u p f m v an d el m b r es en m b st . m b v a. r es en m b a rn u m f m o d d . f m l o w er m io ce n e a rn u m f m b as tr u p f m l o w er m io ce n e 0 cps cps 120 fig. 57. type and reference sections of the vandel member. the type section is the interval from 114 to 102 m in the vandel mark borehole. the reference section is the interval from 100 to 97 m in the grindsted borehole; for legend, see fig. 8, p. 17. odd.: odderup. st.: stauning. va.: vandel mb. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 59 with minor content of mica. heavy minerals are locally very common. the sand is characterised by low-angle crossbedding dipping towards the south-west, and is enriched in heavy minerals (fig. 59). the fine-grained part of the formation is dominated by hummocky cross-stratified sand and homogenous to laminated sand. the odderup formation is characterised by a succession of sand with subordinate clay layers; the odderup formation is differentiated from the arnum formation in being sand-dominated with a sand/mud ratio of at least 75% and a minimum thickness of 5 m. log characteristics. the formation is characterised by low to moderate gamma-ray values (fig. 58); an overall decreasing-upward gamma-ray trend is typical. high gamma-ray values are associated with beds rich in heavy minerals. fossils. marine molluscs as well as dinocysts occur in the south-western sections of the odderup formation (stauning member; piasecki 1980; dybkjær & piasecki 2010). foraminifers reported from coarser-grained (more proximal) intervals (laursen & kristoffersen 1999) may be the result of caving from higher strata. fossil seeds, leaves and wood are abundant in coal beds and lacustrine sands and muds of the terrestrial fasterholt member. depositional environment. the odderup formation was deposited in the lower to upper shoreface and swash zone of a prograding coastal-plain (odderup formation undifferentiated and stauning member). the coals and associated sediments are the deposits of freshwater lakes, lagoonal swamps and mires (fasterholt member; koch 1989). boundaries. the lower boundary is placed where fossiliferous, dark brown, silty clays with subordinate, fine-grained sand layers referred to the arnum formation are overlain by a significant thickness (> 5 m) of grey fine-grained sand (sand: mud > 75%), commonly with a high content of heavy minerals. on the gamma-ray log, the lower boundary may be an abrupt shift to lower values, particularly where the fasterholt member directly overlies the arnum formation. this boundary may be more difficult to locate where the stauning member forms the lowermost odderup formation but the increase in the proportion of sand at this level is generally reflected by a fall in the gamma-ray values (e.g ulfborg borehole, plate 5). the upper boundary is a marked change in lithology from the white, fineto medium-grained sand of the odderup formation to the dark brown, clayey silt of the hodde formation. the boundary is typically sharp but locally is marked by a gravel layer, the base of which defines the boundary. the gamma-ray log typically shows a prominent shift (to higher values) at the boundary. 60 rødding borehole dgu no. 141.1141 store vorslunde borehole dgu no. 104.2325 120 0 150 cps cps 110 100 90 80 70 60 50 40 30 30 40 20 10 0 m.b.s.m.b.s. gr gr o d d er u p f m a . f m fa . l o w er m io ce n e o d d er u p f m a . f m a rn u m f m st au n in g m b st au n in g m b o d d er u p f m h o d d e fm l o w er m io ce n e m id d le m io ce n e 0 40 80 fig. 58. reference sections of the odderup formation; for legend, see fig. 8, p. 17. the primary reference section is the interval from 37 to 1 m in the store vorslunde borehole. the secondary reference section is the composite interval (118–110 m, 90–41 m) in the rødding borehole. a.: arnum. fa.: fasterholt mb. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 60 distribution. the odderup formation is distributed in west, central and southern jylland (fig. 10g). biostratigraphy. the cordosphaeridium cantharellus, exochosphaeridium insigne, cousteaudinium aubryae and labyrinthodinium truncatum dinocyst zones of dybkjær & piasecki (2010) occur in the marine parts of the odderup formation. geological age. the odderup formation is of burdigalien to early langhian (early to earliest middle miocene) age. subdivision. the odderup formation includes the new stauning member and the fasterholt member (koch 1989). stauning member new member history. knudsen et al. (2005) recognised that fine-grained sand layers with a high content of heavy minerals occurred in the arnum formation in a number of boreholes in south and central jylland; these sand layers were informally referred to as the ‘stauning sand’. on gamma-ray logs, the sand beds are characterised by extremely high gamma-ray values. exploration for these heavy mineral sands was intensive during the latter part of the 1990s in the stauning and give areas. name. after the village of stauning (fig. 1) where the member subcrops quaternary deposits at relatively shallow depths. type and reference sections. the type section of the stauning member is defined in the interval from 95 to 76 m (95–76 m md) in the vandel mark borehole (dgu no. 115.1371; 55°42´47.99´´n, 9°10´54.82´´e; fig. 60). the reference section is the intervals from 118 to 110 m (118–111 m md) and 90 to 63 m (90–64 m md) in the rødding borehole (dgu no. 141.1141; fig. 60). thickness. intervals referred to the stauning member commonly range from 10 to 40 m in thickness (e.g. plates 2, 3), but over 100 m has been found in the extreme southern part of the study area, for example in the tinglev borehole (plate 9). lithology. intervals assigned to the stauning member, by definition, have a sand/mud ratio of at least 75% and are more than 5 m thick. the member is typically composed of grey to white, fine-grained sand, with a high content of heavy minerals, intercalated with dark brown, clayey silt (fig. 61). log characteristics. the member typically shows a highly serrated gamma-ray log (e.g. plate 6, stauning borehole) although some sections show more stable low gamma-ray values (e.g. tinglev borehole, plate 9). extremely high gamma-ray readings (e.g. kvong borehole, plate 4; løvlund borehole, plate 7) are found in association with concentrations of heavy minerals. fossils. marine molluscs occur in the stauning member (knudsen 1998) as well as foraminifers and dinocysts (laursen & kristoffersen 1999; dybkjær & piasecki 2010). 61 fig. 59. exposure (isenvad gravel pit) of the odderup formation showing lowangle cross-bedded sand with concentrations of dark heavy minerals; the sand was deposited in the swash zone of a beach. the height of the illustrated section is 0.4 m. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 61 depositional environment. the stauning member was deposited as storm sand layers on the inner shelf, the sands being primarily of storm origin. boundaries.the lower boundary is placed at the base of sanddominated (>75% sand) successions at least 5 m thick, overlying the mud-rich arnum formation. in some wells, this boundary is marked by a general upward decrease in the background gamma-ray values (e.g. hellevad borehole, plate 1; føvling borehole, plate 3) but anomalous exam62 vandel mark borehole dgu no. 115.1371 100 90 80 70 60 50 40 30 m.b.s. rødding borehole dgu no. 141.1141 120 0 150 110 100 90 80 70 60 50 40 m.b.s. grgr o d d er u p f m h . o d d er u p f m o d d er u p f m o d d er u p f m st au n in g m b st au n in g m b st au n in g m b a rn u m f m a r. fm a rn u m f m m id d le m io ce n e l o w er m io ce n e l o w er m io ce n e 0 60 cps cps 120 fig. 60. type and reference sections of the stauning member; for legend, see fig. 8, p. 17. the type section is the interval from 95 to 76 m in the vandel mark borehole. the reference section is the composite interval (118–110 m, 90–63 m) in the rødding borehole. ar.: arnum. h.: hodde fm. 10 cm a b ca b c fig. 61. cores of the stauning member showing homogenous to laminated, grey sand with some intercalated dark brown muds. note the high content of shells in c (arrows), especially in the lower part of the sand beds. a: 15.72 m (base of illustrated core); b: 23.73 m (base); c: 26.94 m (base). cores from a shallow borehole to investigate the heavy mineral potential of stauning member sands; 2 km due west of skjern. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 62 ples are also observed (e.g. rødding borehole, fig. 60), possibly due to the heavy mineral content of the sands. the upper boundary is defined where the fine-grained sand-rich succession is overlain by dark brown, silty clay of the arnum formation or mediumto coarse-grained sand of the odderup formation. where the odderup formation succeeds the stauning member, the upper boundary is commonly reflected by a shift from a dominantly serrated gamma-ray log pattern to a steady and gradually decreasing gamma-ray log pattern (e.g. plates 1, 2). distribution. the stauning member is found in southern, central and western jylland (fig. 10g). biostratigraphy. the cordosphaeridium cantharellus, exocho sphaeridium insigne, cousteaudinium aubryae and labyrinthodinium truncatum dinocyst zones of dybkjær & piasecki (2010) are recognised in the stauning member. geological age. the stauning member is of burdigalian to early langhian (early to earliest middle miocene) age. fasterholt member history. the fasterholt member was defined by koch (1989). brown-coal-bearing layers were mentioned by forchhammer (1835) and brown-coal beds that crop out in the banks of the skjern å (river) were reported by dalgas (1868) and hartz (1909). extensive mining of brown coal occurred during the two world wars and large prospecting programs were carried out in connection with the demands for local energy resources (milthers 1939; milthers 1949). name. after the village of fasterholt (fig. 1). type and reference sections.the formation has previously been exposed in several brown-coal pits in central and western jylland and the fasterholt brown-coal pit (56°00´52.60´´n, 9°06´16.05´´e) is the type locality of koch (1989; figs 1, 62). the member is only exposed today in a small pit at abildå near ørnhøj (fig. 1). the reference section is defined in the store vorslunde borehole (dgu no. 104.2325) from 15 to 13 m (15–13 m md; fig. 62). thickness. the member is c. 8.5 m thick in the type section and is commonly about 10 m thick elsewhere in central jylland (plate 2). it is not recognised in south-west jylland (fig. 10g). lithology. the fasterholt member consists of interbedded sands, clays and brown coals. in the type section, it consists of three sedimentary units, each typically showing a fining-upward trend from a basal sandy lower part passing upward into silty clay and capped by a brown-coal layer (fig. 62). fossils. marine fossils are absent but spores and pollen, fossil seeds, leaves and wood occur abundantly (christensen 1975, 1976; friis 1975, 1979; koch 1977, 1989; koch & friedrich 1970; koch et al. 1973; wagner & koch 1974). depositional environment. the member is interpreted to represent deposition in a terrestrial setting that included lacustrine and mire environments (koch 1989). the con63 fig. 62. type and reference sections of the fasterholt member; for legend, see fig. 8, p. 17. the type section is the fasterholt brown coal pit, north-west of brande; this section is no longer exposed. the log is redrawn from koch (1989). the reference section is the interval from 15 to 13 m in the store vorslunde borehole. a.: arnum fm. fa.: fasterholt mb. fasterholt outcrop 0 1 2 3 4 5 6 7 8 9 10 cl si f mc p sand store vorslunde borehole dgu no. 104.2325 40 30 20 10 0 m.b.s.m gr o d d er u p f m l o w er t o m id d le m io ce n e fa st er h o lt m b o d d e ru p f m a . l o w er t o m id d le m io ce n e f a. 0 40 cps 80 bulletin 22_ gsb191-indhold 04/03/11 12.41 side 63 centration of brown coals in the depocentre of the norwegian–danish basin, particularly adjacent to preexisting faults indicates a structural control on the deposition. boundaries. the lower boundary is sharp, being placed where white sands are overlain by a succession dominated by silty clay and brown coal, with intercalated sands. the lower boundary may be marked by a dense root horizon with tree stumps. on the gamma-ray log, the boundary is characterised by a prominent shift towards high gammaray values. the upper boundary is also sharp, being typically marked by the incoming of the sand-rich upper part of the odderup formation; this lithological change is indicated on the gamma-ray log by a distinct shift to lower readings. where overlain by clay-rich sediments of the arnum formation (e.g. vind borehole, plate 4) or the hodde formation (e.g. fjelstervang borehole, plate 3), the gamma-log values show an abrupt increase. distribution. the fasterholt member is restricted to central jylland (fig. 10g). biostratigraphy. in the absence of marine fossils, the fasterholt member is stratigraphically constrained by the presence of the c. aubryae dinocyst zone below (in the marine odderup or arnum formations) and the l. truncatum dinocyst zone above (in the overlying arnum formation) (dybkjær & piasecki 2010). geological age. due to the absence of marine fossils, the fasterholt member is dated indirectly by the biostratigraphy of the underand overlying marine strata. the age of the fasterholt member is thus constrained to burdigalian to early langhian (early to earliest middle miocene). 64 bulletin 22_ gsb191-indhold 04/03/11 12.41 side 64 history.the succession defined here as the måde group was referred by l.b. rasmussen (1961) to the ‘måde serien’; this encompassed the marine, clay-dominated younger miocene deposits. as described by l.b. rasmussen (1961), the succession is characterised by a basal gravel layer which is overlain by black, mica-rich mud followed by a thin green ish, glaucony-rich clay, grey clay and finally by fineto mediumgrained sand. relative to the north sea litho stratigraphy, the måde group correlates with the nordland group (deegan & scull 1977; hardt et al. 1989; schiøler et al. 2007). name. after a local area west of esbjerg (fig. 1) that was renowned for its brickworks based on upper miocene clays; the last brick factories were closed in the 1970s. type area. the type area of the måde group is south-west jylland. the group is exposed at the gram clay-pit (fig. 1) where both the gram and marbæk formations can be seen. at ørnhøj (lille spåbæk), the hodde and ørnhøj formations are exposed and the marbæk formation crops out in coastal cliffs at sjelborg and marbæk, north-west of esbjerg (fig. 1). the full development of the group is illustrated by the cored borehole sdr. vium (dgu no. 102.948; 51–24 m, fig. 63) and the tinglev borehole (dgu no. 168.1378) from 197 to 50 m (197– 49 m md; fig. 63). thickness. the group is typically about 25 m thick in the western part of jylland, but in southernmost jylland, for example in the tinglev borehole, nearly 150 m has been penetrated (fig. 63; plate 9). lithology. the måde group is dominated by dark brown, organic-rich mud (fig. 64). the lower part is composed of alternating fine-grained sand and silty clay with a basal gravel layer (hodde formation). upwards, the succession becomes more fine-grained with scattered incursions of glaucony. this is succeeded by greenish brown, glauconyrich clay, typically 3 m thick (ørnhøj formation). in the upper part of the glaucony-rich section, goethification of glaucony grains is common (dinesen 1976). this is overlain by a succession of brown clays rich in pyrite that becomes siltier upwards with thin (c. 5 cm thick), finegrained storm sand beds occurring in the upper part. the uppermost måde group consists of fineto mediumgrained sand. log characteristics. the group is characterised by moderate to high gamma-ray readings (fig. 63). extremely high gamma-ray values may be recorded in the lower levels of the group, the upper part showing a gradual decrease in gamma-ray readings (see fig. 63). 65 måde group new group sdr. vium borehole dgu no. 102.948 60 50 40 30 m.b.s. tinglev borehole dgu no. 168.1378 200 0 800 cps 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 m.b.s. grgr m åd e g ro u p r ib e g ro u p a rn u m f m h o d . f m ø . f m g ra m f m m id d le m io ce n e u p p er m io ce n e l . m io . m id d le m io ce n e u p p er m io ce n e m åd e g ro u p r ib e g rp o d d er u p f m h o d d e fm * g ra m f m m ar b æ k fm 20 cps 140 fig. 63. the full development of the måde group is illustrated by the interval from 51 to 24 m in the cored sdr. vium borehole and the interval from 197 to 50 m in the tinglev borehole; for legend, see fig. 8, p. 17. hod.: hodde. l. mio.: lower miocene. ø and *: ørnhøj fm. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 65 fossils. the måde group contains rich and diverse mollusc faunas, crustaceans and vertebrates. shark teeth are common. foraminifers and dinocysts are abundant (see details below in the description of the individual formations). depositional environment. the måde group was deposited on a marine shelf. when the flooding of the land was at its maximum, during the deposition of the glaucony-rich ørnhøj formation and the lower part of the gram formation, the water depth was over 100 m (laursen & kristoffersen 1999). the upper part of the group was deposited in front of a prograding coastline in an offshore to shoreface setting. boundaries. the lower boundary is sharp, being marked by a thin gravel layer separating the white, fine-grained sand of the ribe group from the dark brown mud of the måde group. the upper boundary is a sharp erosional boundary separating mud and fine-grained sand of the måde group from quaternary deposits, the boundary commonly being characterised by a distinct change in lithology and colour of the deposits. distribution. the måde group is restricted to the western and southern part of jylland (fig. 10h) and is found locally around herning and in the brande–give area (fig. 1). geological age. the måde group is of early langhian to latest tortonian (early middle to late miocene) age. subdivision. the måde group is divided into four formations: the hodde, ørnhøj, gram and marbæk formations. hodde formation history. the hodde formation was defined by l.b. rasmussen (1961) from the hodde-1 borehole; it was exposed during the construction (1941–43) of the 66 fig. 65. reference sections of the hodde formation; for legend, see fig. 8, p. 17. the primary reference section is the interval from 51 to 44.9 m in the cored sdr. vium borehole. the secondary reference section is the interval from 50 to 39 m in the føvling borehole. g.: gram fm. hod.: hodde. l. mio.: lower miocene. u.: upper miocene. ør. and ørn.: ørnhøj. sdr. vium borehole dgu no. 102.948 60 60 cps 100 50 40 30 m.b.s. føvling borehole dgu no. 132.1835 60 60 180 cps 50 40 m.b.s. grgr m id d le m io ce n e u p p er m io ce n e ø rn . f m g ra m f m h o d . f m a rn u m f m m id d le m io ce n e u . l . m io . h o d d e fm ø r. fm g . o d d er u p f m marine middle–upper miocene deposits (måde group) continental lower middle miocene deposits (ribe group) continental lower middle miocene deposits (ribe group) fig. 64. the open pit at lille spåbæk, ørnhøj (fig. 1) where the hodde, ørnhøj and gram formations were exposed in the late 1970s. these three formations, together with the marbæk formation, constitute the måde group. the cliff is c. 10 m high. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 66 karlsgårde channel, near hodde, but this exposure does not exist today. name. after the village of hodde in south-west jylland (fig. 1). type and reference sections. the type section was defined by l.b. rasmussen (1961) as the interval from 23.4 to 13.8 m in the hodde-1 borehole (dgu no. 113.33; 55°41´04.11´´n, 8°40´14.27´´e). the formation is exposed at lille spåbæk near ørnhøj, south of holstebro (fig. 1). the primary reference section is the interval from 51 to 44.90 m in the cored borehole at sdr. vium (dgu no. 102.948; fig. 65). a secondary reference section is the føvling borehole (dgu no. 132.1835) from 50 to 39 m (fig. 65). thickness. the formation is 9.6 m thick (23.4–13.8 m) in the type section and is typically 5–10 m thick where present, but thickens in southernmost jylland; more than 40 m was penetrated in the rømø borehole (plate 9). lithology. the hodde formation consists of dark brown, organic-rich, bioturbated silty clay with thin sand lenses (figs 66, 67); the pyrite content is high. the basal part of the formation is composed of a thin gravel layer. in the upper part of the formation, laminated, silty clay is common and glaucony may occur. trace fossils are common in the hodde formation (asgaard & bromley 1974). log characteristics. the formation is typified by moderate to high gamma-ray values (fig. 65); a gradual upward increase in gamma-ray response is characteristic. locally, the upper part shows low gamma-ray readings, for example in the føvling borehole (plate 8). fossils. the hodde formation typically contains a limited fauna of marine molluscs (l.b. rasmussen 1966) but a 67 fig. 66. the upper part of the hodde formation at lille spåbæk, ørnhøj, dominated by dark brown silty clay. fig. 67. close-up of the hodde formation at lille spåbæk, ørnhøj. the hodde formation is composed of dark brown silty clay; the yellowish stripes are due to weathering of pyrite. the illustrated section is 0.5 m high. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 67 richer fauna occurs locally in shell-beds associated with the basal gravel bed. marine microfossils, such as foraminifers and dinocysts, occur abundantly (laursen & kristoffersen 1999; piasecki 1980, 2005; dybkjær & piasecki 2010). depositional environment. the depositional environment is interpreted as fully marine (l.b. rasmussen 1961). the basal coarse-grained transgressive lag indicates deposition on a marine shoreface during the initial transgressive phase. the increase in glaucony in the upper part indicates a near cessation of sediment influx to this part of the north sea in the serravallian. boundaries. there is a marked change in lithology from the white, fineto medium-grained sand of the odderup formation to the overlying dark brown, clayey silt of the hodde formation. the boundary is sharp and is commonly characterised by a gravel layer, the base of which (where present) defines the boundary. the gamma-ray log shows a prominent shift to high values at the lower boundary. the upper boundary is defined by an abrupt change from dark brown, clayey silt of the hodde formation to greenish brown clay of the ørnhøj formation. on the gamma-ray log, this is reflected by a distinct shift towards higher gamma-ray values. distribution. the hodde formation is recognised in southern and western jylland (fig. 10h). the formation occurs locally as far east as bording and give in central jylland in depressions associated with salt structures. biostratigraphy. the upper part of the labyrinthodinium truncatum dinocyst zone and the unipontidinium aquaeductum dinocyst zone (dybkjær & piasecki 2010) are recorded in the hodde formation. geological age. the hodde formation is of early langhian to mid-serravallian (middle miocene) age. ørnhøj formation new formation history. formerly referred to as the ‘glauconitic clay member’ of the lower gram formation of previous usage (l.b. rasmussen 1956, 1961). name. after the village of ørnhøj (fig. 1) where the formation is still exposed in some of the old brown-coal pits in the neighbourhood. type and reference sections. the formation is partly exposed at lille spåbæk, west of ørnhøj (figs 1, 64). the type section is the interval from 44.90 to 40 m in the cored borehole at sdr. vium (dgu no. 102.948; 55°49´04.02´´n, 8°24´46.52´´e; fig. 68). the reference section is the inter68 fig. 68. the type and reference sections of the ørnhøj formation; for legend, see fig. 8, p. 17. the type section is the interval from 44.9 to 40 m in the cored sdr. vium borehole. the reference section is the interval from 39 to 36 m in the føvling borehole. g: gram fm. ho.: hodde. l. mio.: lower miocene. m.: middle. odd.: odderup. ø.: ørnhøj fm. *: upper miocene. føvling borehole dgu no. 132.1835 sdr. vium borehole dgu no. 102.948 60 60 cps 240 50 40 30 20 10 0 m.b.s. gr m.b.s. gr 45 120 cps 160 44 43 42 41 40 39 38 glaucony cl si f mc p sand ø rn h ø j fm ø . g h o . f m o d d . f m g ra m f m m id d le m io ce n e m . m io ce n e * q u at er n ar y l . m io . u p p er m io ce n e bulletin 22_ gsb191-indhold 04/03/11 12.41 side 68 val from 39 to 36 m (39–36 m md) in the føvling borehole (dgu no. 132.1835; fig. 68). thickness. the formation is 4–5 m thick in the type and reference boreholes, but in general it rarely exceeds more than 2 m in thickness (plates 4, 8, 9). lithology. the ørnhøj formation is composed of green and brown clay (fig. 69). high concentrations of green glaucony pellets of fine sand grade occur commonly. in the upper part of the formation, goethification of glaucony is common. log characteristics. the formation is characterised by high gamma-ray values. fossils. the ørnhøj formation is barren of macroand microscopic calcareous fossils but a diverse assemblage of dinocysts is present (piasecki 1980, 2005; dybkjær & piasecki 2010). depositional environment. the ørnhøj formation was deposited in a fully marine, sediment-starved depositional setting that favoured the formation of glaucony. the water depth was probably more than 100 m, based on the estimates of water depth during deposition of the gram formation (see below). the ørnhøj formation represents the most widespread transgression during the miocene (e.s. rasmussen 2004b; knox et al. 2010). the goethification of glaucony in the upper part is interpreted as a result of a sea-level fall (dinesen 1976; eder et al. 2007) with associated wave action at the sea floor. concentration of glaucony in depositional bars at ørnhøj (j. frederiksen, personal communication 2009) supports the interpretation of wave action at the sea floor. boundaries. the lower boundary is characterised by an abrupt change from the dark brown, clayey silts of the hodde formation to greenish brown clays of the ørnhøj formation (fig. 69). the gamma-ray log shows a prominent shift in gamma-ray response towards high values. the upper boundary is defined by the change from greenish brown or brown, glaucony-rich clay to dark brown clay. at the boundary there is an abrupt change from glaucony-impregnated pellets and shells to pyritised pellets. on the gamma-ray log, the upper boundary is defined at a decrease in gamma-ray values; locally a very prominent decrease is observed, for example in the stensig borehole (plate 4). distribution. the ørnhøj formation is recognised in southern and western jylland. the formation is locally recognised in the subsurface as far east as bording and give in central jylland, where it occurs in depressions associated with salt structures (fig. 10h). biostratigraphy. the achomosphaera andalousiense and gramocysta verricula dinocyst zones of dybkjær & piasecki (2010) are recorded in the ørnhøj formation. geological age. the ørnhøj formation is of late serravallian (late middle miocene) age. 69 ørnhøj fm hodde fm ørnhøj fm hodde fm fig. 69. the ørnhøj formation at lille spåbæk, ørnhøj. the lower boundary with the hodde formation beneath is seen in the lower part of the section. knife for scale, c. 20 cm long. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 69 gram formation redefined formation history. the gram formation was defined by l.b. rasmussen (1956). in the original definition of the gram formation, three members were recognised: the glauconite clay, gram clay and gram sand members (l.b. rasmussen 1956). the glauconite clay member of previous usage is herein redefined as the new ørnhøj formation and the gram sand member as the marbæk formation; the redefined gram formation thus equates to the gram clay member of l.b. rasmussen (1956). name. after the town of gram (fig. 1). type and reference sections. the type section is at the disused pit of the gram brickworks (55°18´24.90´´n, 9°03´31.26´´e; fig. 1), now the midtsønderjyllands museum of gram, where a 13.1 m thick section of the gram formation is exposed (figs 70, 71). the reference section is the interval from 40 to 24 m in the cored borehole sdr. vium (dgu no. 102.948; fig. 71). thickness. a 13.1 m section is seen at the type section, but neither the base nor the top is exposed. in the reference section, the formation is about 16 m thick. the formation thickens south-westward and 105 m was penetrated in the tinglev borehole (plate 1). lithology. the gram formation consists of dark brown clay, which becomes more silty upwards. in the upper part, a few, fine-grained, wave-rippled sand beds, c. 5 cm thick, are intercalated with the clays (figs 70, 71). siderite concretions are common in the lower part of the formation. pyrite is common both as pyritised pellets and in trace fossils; the latter include common trichichnus ispp. (rasmussen & larsen 1989; bromley 1996). log characteristics. the formation is characterised by moderate gamma-ray values (fig. 71). the log pattern is serrated and shows a general decreasing-upward trend in gamma-ray values through the succession (fig. 63). fossils. the gram formation is characterised by abundant and diverse mollusc faunas, in association with marine vertebrates (whales and sharks) and crustaceans (crabs), the latter in concretionary nodules (l.b. rasmussen 1966, 1968; bendix-almgreen 1983; hoch 2008; schnetler 2005; steeman, 2009). foraminifers and dinocysts are abundant (laursen & kristoffersen 1999; piasecki 1980, 2005). depositional environment. the gram formation was deposited in a fully marine environment with water depths of more than 100 m (laursen & kristoffersen 1999; c. morigi, personal communication 2010). the incoming of storm beds in the upper part is interpreted to reflect progradation of the shoreline (rasmussen & larsen 1989). boundaries. the lower boundary is defined at the change from greenish brown or brown, glaucony-rich clay to dark brown clay, associated with an abrupt change from glaucony-impregnated pellets and shells to pyritised pellets. on the gamma-ray log, this is reflected by a decrease in gamma-ray values. 70 fig. 70. fine-grained, partly bioturbated sand interbeds in the upper part of the gram formation, gram clay pit. the sand beds are commonly wave-rippled. the illustrated section is 0.30 m high. facing page: fig. 71. type and reference sections of the gram formation; for legend, see fig. 8, p. 17. the type section is at the gram clay pit near gram, where 13.1 m of the formation is exposed. the reference section is the interval from 40 to 24 m in the cored sdr. vium borehole. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 70 71 u p p er m io ce n e g ra m f m u p p er m io ce n e g ra m f m ø rn h ø j fm u p p er m io ce n e u p p er m io ce n e u p p er m io ce n e m ar b æ k fm g ra m f m g ra m f m 0 1 2 3 4 5 6 7 8 9 10 m m.b.s. gr 11 10 12 13 1 2 41 40 39 38 37 36 35 34 33 32 31 30 31 80 100 cps 29 28 27 26 25 24 cl si f mc p sand cl si f mc p sand cl si f mc p sand 110 140 cps cl si f mc p sand gram outcrop sdr. vium borehole dgu no. 102.948 covered interval (< 2m) 0 bulletin 22_ gsb191-indhold 04/03/11 12.41 side 71 72 the upper boundary is placed where interbedded clay and thin sand layers are succeeded by amalgamated sand beds. on the gamma-ray log, the upper boundary is identified by a marked shift to consistently low gamma-ray values. distribution. the gram formation is recognised in the subsurface of southern and western jylland (fig. 10h). the formation occurs locally as far east as bording and give in central jylland in depressions associated with salt structures. biostratigraphy. the amiculasphaera umbracula and hystrichosphaeropsis obscura dinocyst zones of dybkjær & piasecki (2010) are recorded in the gram formation. geological age. the gram formation is of tortonian (late miocene) age. marbæk formation new formation history. sands exposed in the cliffs at sjelborg and marbæk, north-west of esbjerg (fig. 1), and sandy sediments in the upper part of the sæd borehole (dgu no. 167.445) were tentatively referred to the pliocene by jørgensen (1945). new studies of these sections (piasecki et al. 2003), however, indicated that these deposits are tortonian in age. the sand was informally named the gram sand member (gram formation of previous usage) by l.b. rasmussen (1956). name. after the coastal cliff at marbæk, north-west of esbjerg (fig. 1). type and reference sections. the type section is the exposure at marbæk cliff (55°32´56.49´´n, 8°18´57.49´´e; figs 1, 72). the reference section is the interval from 62 to 50 m (62–49 m md) in the tinglev borehole (dgu no.168.1378; fig. 73). thickness. the marbæk formation is c. 16 m thick in the marbæk cliff (fig. 72); neither the base nor the top is exposed. in the pit at the gram brickworks, 1.5 m of the formation is exposed in the bank of a stream (fig. 71). in marbæk outcrop 0 1 2 3 4 5 6 7 8 9 10 10 11 12 13 14 15 cl si f mc p sand cl si f mc p sand ? ? ? ? ? m u p p er m io ce n e m ar b æ k fm u p p er m io ce n e m ar b æ k fm fig. 72. type section of the marbæk formation in the coastal cliff at marbæk, north-west of esbjerg, where c. 16 m of the formation are exposed; for legend, see fig. 8, p. 17. iron-stained fractures are conspicuous at 7–9 m in this section. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 72 73 the tinglev borehole, the formation is 10 m thick (fig. 73, plate 1), though the top is an unconformity with quaternary sediments. lithology. in the type section, the formation is dominated by white, often reddish, fineto medium-grained, mica-rich sand with a few thin intercalated coarse-grained sand or gravel layers and, in the lower part, subordinate silt-rich intervals (figs 72, 74). the sand beds show parallel lamination with subordinate cross-bedding; hummocky crossstratification is common (fig. 75). a silt-rich interval shows double clay layers. the uppermost white sand at sjelborg consists of homogenous sand capped by wave-ripples (fig. 76). the pyrite content is very high in the marbæk formation (olivarius 2009) and the distinctive red colour of the succession at the marbæk outcrop is due to oxidation of the pyrite. fossils. rare, poorly preserved molluscs have been found in the marbæk formation (jørgensen 1945). dinocysts occur in the lower part of the formation but become scarce upwards (piasecki et al. 2003). depositional environment. the formation was deposited in a storm-dominated environment within the upper and tinglev borehole dgu no. 168.1378 90 80 70 60 50 m.b.s. gr u p p e r m io ce n e m ar b æ k f m g ra m f m 0 800 cps fig. 73. reference section for the marbæk formation is the interval from 62 to 50 m in the tinglev borehole; for legend, see fig. 8, p. 17. fig. 74. oblique view of the marbæk formation at marbæk, north-west of esbjerg. two persons (upper right) for scale. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 73 74 lower shoreface. double clay layers indicate some tidal influence. boundaries. the lower boundary is defined where alternating thin clay and sand layers are overlain by amalgamated sand beds; this boundary is not observed at outcrop. on the gamma-ray log, this boundary is identified by a marked shift to steady low gamma-ray values. the upper boundary is placed at a distinct, erosional unconformity separating the mica-rich sands from tills and yellowish, coarse-grained sands and gravels of quaternary age. distribution. the formation is limited to the far west and south of jylland. biostratigraphy. the hystrichosphaeropsis obscura dinocyst zone of dybkjær & piasecki (2010) is recorded in the lower part of the marbæk formation. geological age. the lower part of the marbæk formation is of tortonian (late miocene) age, equivalent to the uppermost part of the gram formation. the absence of fossils in the upper levels of the formation precludes precise dating of this part. fig. 75. hummocky cross-stratified sand of the marbæk formation at marbæk; the illustrated section is 50 cm high. fig. 76. although deformed by glacial tectonics, the marbæk formation sands display homogenous and wave-rippled facies typical of upper shoreface deposits. the illustrated section is 0.4 m high; sjelborg. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 74 75 the overall stratigraphic architecture of the miocene succession is best revealed by integration of seismic data with outcrop and borehole data. thus, in the grid of correlation panels presented here (plates 1–9), the borehole and outcrop data provide the critical, stratigraphic constraints whilst the sedimentary architecture between wells is based in large part on the seismic data. inspection of seismic sections (figs 77–79) reveals that the lower part of the miocene succession is composed of two discrete, progradational successions (fig. 77). the first succession includes the vejle fjord and billund formations and the second succession includes the klintinghoved and bastrup formations. these packages are often characterised by a seismic reflection pattern that shows both oblique–parallel and sigmoidal clinoforms. the height of clinoforms ranges between 60 m and 100 m, and dips of the clinoforms commonly vary between 3° and 10° (fig. 77). clinoformal packages may alternate with units of more or less transparent seismic character (fig. 78). this part of the succession is interpreted to represent prograding delta lobes with alternating sand-rich and mud-rich units (e.s. rasmussen et al. 2007; hansen & rasmussen 2008; e.s. rasmussen 2009b; fig. 80, plates 1–9). on top of each prograding unit, erosional valleys and channels occur and some channels are characterised by having a shingled reflection pattern (e.s. rasmusssen 2009b). these features were formed by incision and are commonly filled with sand. the shingled reflection pattern is interpreted to represent point bars of meandering river systems (e.s. rasmussen et al. 2007; e.s. rasmussen 2009b). in between these delta lobes, seismic reflectors are parallel, commonly of low amplitude (fig. 78); this seismic character is considered to reflect the presence of mud-dominated inter-lobe deposits (hansen & rasmussen 2008). in northern and central jylland, a successive southward migration of delta lobes can be demonstrated (e.g. plate 2), defining an ascending shoreline trajectory (fig. 78) indicating progradation during rising sea level (e.g. helland-hansen & gjelberg 1994). in the northern part of the study area, the lower miocene is dominated by a parallel to subparallel reflection pattern capping the clinoforms (fig. 79). boreholes penetrating this part of the succession indicate alternating mudand sand-rich units (fig. 79). gravel pits and outcrops around silkeborg indicate a dominance of braided fluvial systems stratigraphic architecture 0 100 200 t w t ( m se c) 300 (offset 4 km) northsouth store vorslunde billundalmstok 2 km 5 0 m top delta lobe top bastrup top billund/vejle intra odderup base miocene internal reflector/surface clinoforms incised valleys base quaternary bastrup delta complexbastrup delta complex resen mbresen mb addit mbaddit mb arnum fmarnum fm arnum fm arnum fm arnum fmarnum fm arnum fm billund delta complexbillund delta complex odderup coastal plain odderup coastal plain odderup fmodderup fm vejle fjord fmvejle fjord fm vejle fjord fm vejle fjord fm klintinghoved fm klintinghoved fm fig. 77. s–n-trending seismic section tieing the almstok, billund and store vorslunde boreholes (for location, see fig. 1). sand-rich delta lobes characterised by an oblique–parallel reflection pattern are indicated in yellow. note the alternation of these sand-rich delta deposits and more clay-rich inter-lobe deposits, a characteristic feature of the vejle fjord–billund formations and the klintinghoved–bastrup formations. the upper part of the section is dominated by a parallel to subparallel reflection pattern which is characteristic of the arnum–odderup formations and indicates a change in sedimentation style. seismic data courtesy of cowi a/s. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 75 76 100 200 300 400 500 hammerum 200 150 100 50 0 isenvad (offset 300 m) 150 100 50 billund fm billundfmaddit mb addit mb klintinghoved fm resen mbresen mb billund fm billundfm vejle fjord fm vejle fjord fm addit mb addit mb klintinghoved fm resen mbresen mb odderup fmodderup fm bastrup fmbastrup fm eastwest 1 km t w t ( m se c) 1 km resen mb bastrup fm delta lobe delta lobe billund fm resen mb bastrup fm delta lobe delta lobedelta lobe billund fmbillund fm addit mbaddit mb addit mbaddit mb delta lobe billund fm vejle fjord fmvejle fjord fm northsouth 5 0 m klintinghoved fm klintinghoved fm fig. 79. w–e-striking seismic section at ikast (for location, see fig. 1). the section shows a cross-section of the billund and bastrup delta systems as indicated by dipping reflectors both towards the west and east. in this area, the fluvial systems of the addit and resen members, shown in red, are particularly well developed. seismic courtesy of rambøll a/s; for legend, see fig. 77). fig. 78. detailed seismic section of the store vorslunde area (for location, see fig. 1). sand-rich parts of the deltas are indicated in yellow, sand-rich fluvial deposits of the addit and resen members are shown in red. seismic data courtesy of cowi a/s; for legend see fig. 77. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 76 77 in this area (hansen 1985; hansen 1995; jesse 1995; e.s. rasmussen et al. 2007), and based on subsurface data such braided fluvial systems dominated in northern jylland and parts of central jylland from addit to hammerum (plate 6). in a narrow nw–se-striking belt across central jylland, the seismic data show an oblique–parallel reflection pattern. this represents progradation during falling sea level (hansen & 2008, e.s. rasmussen 2009b). in southern jylland, there is a tendency towards a combined aggradational– progradational stacking pattern (plate 1), reflecting progradation during rising sea level. above these two progradational units of the billund/vejle fjord system and the klintinghoved/bastrup system (i.e. above the top bastrup reflector in figs 77–79), a parallel to subparallel reflection pattern dominates the miocene succession (fig. 77); this correlates with the arnum and odderup formations. the change in seismic character indicates a change in depositional environment from prograding ‘gilbert-type’ deltas to aggrading shelf and coastal plain deposits. this is illustrated on correlation panels by progressive outbuilding of the odderup formation towards the south-west contemporaneously with the accumulation of the marine arnum formation (fig. 80; plate 4). bulletin 22_ gsb191-indhold 04/03/11 12.41 side 77 78 78 billund dgu no. 114.1857 estrup dgu no. 132.1838 vorbasse dgu no. 123.1167 rødding dgu no. 141.1141 almstok dgu no. 114.1858 sto dg løgumkloster dgu no. 159.739 odderup fm arnum fm arnum fm arnum fm bastrup fm bastrup fm billund fm odderup fm odderup fm arnum fm vejle fjord fm klintinghoved fm brejning fm stauning mb stauning mb stauning mb stauning mb vandel mb stauning mb stauning mb resen mbresen mb resen resen mbbastrup fm hodde fm ørnhøj fmgram fm odderup fm addit mb addit mb billund fm klintinghoved fm fasterholt mb resen mb resen mb resen mb arnum fm bastrup fm bastrup fm bastrup fm billund fm vejle fjord fm klintinghoved fm isenvad dgu no. 86.2056 a d stauning dgu no. 93.1125 b west assing mølleby dgu no. 94.2821 arnum fm bastrup fm bastrup fm arnum fm billund fm vejle fjord fm stauning mb kolding fjord mb gram fm hodde fm ørnhøj fm stensig dgu no. 93.1062 hammerum dgu no. 85.2429 hjøllund dgu no. 86.2118 50 m ? a south b a bulletin 22_ gsb191-indhold 04/03/11 12.41 side 78 79 resen dgu no. 65.1643 sunds dgu no. 85.2452 billund dgu no. 114.1857 almstok dgu no. 114.1858 store vorslunde dgu no. 104.2325 hammerum dgu no. 85.2429 fasterholt dgu no. 95.2730 klyngholt bastrup fm bastrup fm billund fm billund fm billund fm brejning fm vejle fjord fm odderup fm odderup fmodderup fm arnum fm stauning mb el mb stauning mb stauning mb n mb resen mb resen mb resen mb addit mb fasterholt mb fasterholt mb hodde fm ørnhøj fm gram fm addit mb billund fm billund fm billund fmbillund fm hoved fm vejle fjord fm arnum fm vejle fjord fm isenvad dgu no. 86.2056 addit mark dgu no. 97.928 addit gravel pit dgu no. 97.1000 morsholt dgu no. 108.148 east hjøllund dgu no. 86.2118 brejning fm depositional environments sand (marine) sand/gravel (continental) clay (marine) clay (lagoonal) clay (continental) coal 50 m base quaternary top bastrup fm / klintinghoved fm top billund fm / vejle fjord fm main boundaries top brejning fm top paleocene clay top odderup fm / arnum fm north b a fig. 80. two typical correlation panels showing the overall architecture of the miocene succession in jylland. a: s–n-trending section from løgumkloster to resen. note the change in the depositional style from the forestepping delta lobes of the markedly progradational billund and bastrup formations to the more regular, aggrading – weakly prograding system of the odderup formation. b: w–e-striking section from stauning to morsholt. note that the main delta lobes pinch out both to the east and to the west. the two correlation panels are also shown in plates 2 and 6, together with detailed borehole logs. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 79 the late oligocene was characterised by a warm climate and thus a period with a high sea level (utscher et al. 2000, 2009; zachos et al. 2001; miller et al. 2005; larsson et al. 2006; larsson-lindgren 2009). the north sea was located in the northern westerly wind belt (galloway 2002) and consequently the north-eastern part of this sea, which covered present-day denmark, was dominated by wave processes due to the long fetch across the north sea (fig. 2). most of present-day denmark was covered by the sea in the late oligocene and the deposition of the fully marine brejning formation took place. there is no evidence for the northern position of the coastline at this time, but structural elements such as the sorgenfrei–tornquist zone or the fennoscandian shield were probably important features in controlling the location and trend of the shoreline; the position of the coastline is conservatively placed in the fringe area of the fennoscandian shield (fig. 81a). locally in northern jylland, the diatomite of the sydklint member was formed associated with submarine exposure of eocene diatomites. climatic cooling and initial uplift of the 80 palaeogeography 100 km 100 km a b fig. 81. a: palaeogeographic reconstruction of the latest late oligocene (brejning formation). the exact location of the shoreline is uncertain, but most of present-day jylland was submerged at that time. water depth in northern jylland was over 200 m and extensive formation of glaucony indicates some distance to the shoreline. b: palaeogeographic reconstruction of the earliest early miocene (earliest aquitanian; vejle fjord formation). due to early miocene inversion (reactivation) of the ringkøbing–fyn high and salt structures, a barrier formed between the eastern part of the norwegian–danish basin and the north sea basin. this resulted in brackish water conditions north-east of the ringkøbing–fyn high. small spit systems developed east of these structures. the degradation of these spit systems during the early miocene transgression resulted in the formation of the skansebakke member. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 80 norwegian–danish basin at the end of the oligocene (e.s. rasmussen 2009a) resulted in a fall in relative sea level at the oligocene–miocene transition. this led to deposition of the øksenrade member which was deposited in shallow water on the ringkøbing-fyn high. at dykær, subaerial conditions prevailed for a period (rasmussen & dybkjær 2005). the transition from the oligocene to the miocene was characterised by a short, but marked sea-level fall associated with ice cap growth on antarctica (miller et al. 1996). coincident with this, inversion of the norwegian–danish basin and reactivation of the sorgenfrei–tornquist zone and the ringkøbing–fyn high commenced (e.s. rasmussen 2009a). this resulted in a marked change in the depositional regime in the eastern north sea basin from deposition of dominantly fully marine, clay-rich sediments at the basin floor and toe-of-shelf slope, to sedimentation of coarse-grained, sand-rich, shallow marine, deltaic deposits (larsen & dinesen 1959; l.b. rasmussen 1961; spjeldnæs 1975; friis et al. 1998; michelsen et al. 1998; e.s. rasmussen 1996, 2004a, b). during the earliest miocene, the palaeogeography was controlled by structural highs and lows (fig. 81b). elevated parts of the ringkøbing–fyn high formed a barrier across present-day southern jylland. salt diapirs within the norwegian–danish basin acted as cores of minor islands. a large silled basin formed north of the ringkøbing–fyn high where brackish water conditions prevailed. during 81 100 km 100 km a b fig. 82. a: palaeogeographic reconstruction of the early miocene (aquitanian; billund and vejle fjord formations). the sea level continued to rise during this phase and flooded the ringkøbing–fyn high. due to high sediment supply to the basin, however, the shoreline prograded southward. this favoured the formation of spit/barrier complexes south-east of the main delta lobes; represented by the hvidbjerg member of the billund formation. the river system during the early miocene was dominantly braided in character. b: palaeogeographic reconstruction of the early miocene (late aquitanian; billund formation). during this period, relative sea level fell and progradation of the shoreline is reflected by amalgamation of beach ridges along the coast. distinct incision and formation of broad valleys commenced at the same time. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 81 the lowstand of sea level, sands were transported along the structures and deposited as spits and barrier islands east of the structures (fig. 81b). in this brackish water basin, the lower part of the vejle fjord formation was deposited. a subsequent rise of sea level in the early aquitanian resulted in flooding of the ringkøbing–fyn high. this led to degradation of the barrier complexes and deposition of the skansebakke member. at the time of maximum flooding, the shoreline withdrew to a position north of århus in the east and near thisted in the north-west. the high sediment supply to the north sea basin, however, resulted in progradation of sand-rich delta complexes from the north and north-east, as recorded by the billund formation (fig. 82a). the sediments were probably conveyed through three major river systems (olivarius 2009). the western river was probably connected to the setesdal valley in presentday norway and was the source for the sediments deposited in the delta located off the present west coast of denmark, the so-called ringkøbing lobe (hansen & rasmussen 2008). the central river was sourced from the north, probably from the southern part of present-day norway and the northern part of present-day central western sweden. the eastern river drained the area covered by the present-day central sweden. the central and eastern river system merged in central jylland and resulted in the deposition of the brande lobe of the billund formation (hansen & rasmussen 2008). the river systems were braided and their deposits constitute the addit member. on entering the 82 fig. 83. a: palaeogeographic reconstruction of the early miocene (early burdigalian; klintinghoved formation and kolding fjord member). global climatic warming resulted in a relative rise in sea level. the shoreline was characterised by estuaries and associated barrier complexes and the accumulation of braided fluvial deposits in incised valleys. b: palaeogeographic reconstruction of the early miocene (early burdigalian; klintinghoved formation). during the most widespread flooding in the early burdigalian, most of western and central jylland was covered by the sea and the clay-rich klintinghoved formation was deposited. 100 km 100 km a b bulletin 22_ gsb191-indhold 04/03/11 12.41 side 82 sea, the sands were deposited in wave-dominated deltas (rasmussen & dybkjær 2005; hansen & rasmussen 2008). some of the sand at the delta mouth was reworked and transported eastward by longshore currents to be deposited as spit and barrier complexes of the hvidbjerg member (fig. 82a). development of lagoonal environments was common during this time. due to a global climatic deterioration in the late aquitanian (zachos et al. 2001), sea level began to fall and the delta complexes were forced south-westward; deposition of a coastline characterised by amalgamated beach ridges took place (fig. 82b). a resumed transgression occurred at the beginning of the burdigalian. this transgression was the result of a global warming (zachos et al. 2001). according to t. utescher (personal communication 2008), the average air temperature rose 2°c. widespread barrier-island complexes formed east of the main delta systems due to strong erosion of the main delta and eastward transport of erosional materials. these barrier-island complexes correspond to the kolding fjord member (fig. 83a). during maximum flooding of the sea and associated with the progradation of the succeeding delta complex, mud was deposited in the danish area. this constitutes the klintinghoved formation (fig. 83b). the succeeding delta complex, the bastrup formation, prograded south-westward (fig. 84a) and periodically this progradation occurred during a sea-level fall. the coastline was dominated by beach ridges (fig. 84b). in midburdigalian times, the delta complexes of the bastrup 83 fig. 84. a: palaeogeographic reconstruction of the early miocene (early burdigalian; bastrup formation). progradation occurred during rising relative sea level which formed optimal conditions for a shoreline dominated by lagoons and barrier islands. the fluvial system was dominated by meandering river systems. b: palaeogeographic reconstruction of the early miocene (burdigalian; bastrup formation). during this period, relative sea level fell and the prograding shoreline was characterised by amalgamation of beach ridges parallel to the coast. distinct incision on land and formation of broad valleys commenced at the same time. 100 km 100 km a b bulletin 22_ gsb191-indhold 04/03/11 12.41 side 83 formation reached the southern part of denmark (fig. 85a). during sedimentation of the bastrup formation, the fluvial regime changed character, to be dominated by meandering rivers, especially in the latter phase of progradation. sand deposits of these rivers are referred to the resen member. a distinct global climatic warming, the ‘mid miocene climatic optimum’, occurred at the end of the early miocene (late burdigalian) (zachos et al. 2001). this resulted in a sea-level rise and renewed transgression. the mud laid down during this transgression and in front of the succeeding prograding coastline is represented by the arnum formation (fig. 85b). at the time of maximum flooding, the coastline was located across the northern part of present-day north-west jylland and continued south-eastwards through central jylland. despite a subtropical climate (friis 1975), and hence globally high sea levels during the latest part of early and early middle miocene, progradation resumed (fig. 86a). this was due to tectonism and uplift of the hinterland and consequently increased sediment supply to the north sea basin (e.s. rasmussen 2004b). sand of the prograding coastline is represented by the odderup formation (fig. 86b). fine-grained storm-sand layers deposited in front of the coastline are included in the stauning member. as a consequence of the prograding coastline of the odderup formation and coincident rising sea level due to the warmer climate, conditions for brown coal formation were optimal. a preliminary study (t. 84 fig. 85. a: palaeogeographic reconstruction of the early miocene (burdigalian; bastrup formation) when most of jylland was land; progradation took place during rising sea level. b: palaeogeographic reconstruction of the late early miocene (late burdigalian; arnum formation). the shoreline was located across the northern part of jylland at the time of maximum flooding. 100 km 100 km a b bulletin 22_ gsb191-indhold 04/03/11 12.41 side 84 utescher, personal communication 2009) indicated that widespread coal formation was also associated with increased precipitation in the area. the coal was formed on the coastal plain especially adjacent to pre-existing faults (koch 1989) and predominantly north of the ringkøbing–fyn high. these widespread coal layers are referred to the fasterholt member (koch 1989). due to the overall rising sea level during the mid miocene climatic optimum, and partly also due to ‘auto retreat’ (see muto & steel 2002), a major transgression occurred in the middle langhian (early middle miocene). the transgression was further amplified by increased subsidence of the north sea basin during the middle and late miocene (koch 1989; michelsen et al 1998; clausen et al. 1999; e.s. rasmussen 2004b; e.s. rasmussen et al. 2005). mud-rich sediments of the hodde formation were deposited during this transgression. there is no evidence of the formation of barrier-island complexes during the transgression which suggests a very rapid flooding of the low relief landscape represented by the coal-rich odderup formation. despite major global climatic deterioration (cooling) in the early serravallian (middle miocene; zachos et al. 2001), flooding of this part of the north sea basin continued as a consequence of the accelerating subsidence of the basin. during the most widespread flooding of the area, glaucony-rich sedi ments of the ørnhøj formation accumulated, indicating a long distance to the coastline. the location of the coastline during this maximum transgression is uncertain. 85 fig. 86. a: palaeogeographic reconstruction of the late early miocene (late burdigalian; odderup formation). at this time, the climate became subtropical and the global sea level continued to rise. high sediment supply, however, forced the shoreline to prograde. these conditions favoured formation of lagoons and swamp lakes which were optimal for the formation of coal-rich deposits. b: palaeogeographic reconstruction of the early middle miocene (early langhian; odderup formation). during the maximum regression of the shoreline, most of jylland was land and only the south-western part was submerged. lagoonal and swamp conditions prevailed north of the ringkøbing–fyn high, probably favoured by increased subsidence in this area. 100 km 100 km a b bulletin 22_ gsb191-indhold 04/03/11 12.41 side 85 boreholes in central jylland do not indicate any influx of coarse-grained siliciclastic deposits, so at a minimum, the coastline was displaced to the southern boundary of the fennoscandian shield (fig. 87a). parts of the shield may, however, have been flooded during the highest rate of relative sea-level rise. coincident with the subsidence of the north sea basin, the norwegian mainland was uplifted (løseth & henriksen 2005; rundberg & eidvin 2005; eidvin & rundberg 2007; e.s. rasmussen et al. 2008). this resulted in enhanced sediment supply to the basin where progradation took place. mud of the gram formation was deposited in an open shelf environment (fig. 87b). thin storm-sand layers are intercalated in the upper gram for mation indicating an approaching coastline (rasmussen & larsen 1989). near the end of the tortonian (late mio cene), shoreface sediments of the marbæk formation were deposited in the central–western part of present-day denmark (fig. 88). progradation of the coastline continued through the late miocene and a delta/coastline was formed in the central part of the central graben area (rasmussen 2005; møller et al. 2009). the termination of the miocene was characterised by a sea-level fall of c. 90 m, which is indicated by deep incision of equivalent strata in the offshore cenozoic record of denmark (møller et al. 2009). 86 fig. 87. a: palaeogeographic reconstruction of the middle miocene (serravallian; ørnhøj formation). despite climatic deterioration in the middle miocene, most of jylland was flooded and the shoreline was located in the northern part of jylland. due to very low sedimentation rates, optimal conditions existed for the formation of glaucony. b: palaeogeographic reconstruction of the late miocene (tortonian; gram formation). in the latest part of the miocene, uplift of scandinavia and the alpine mountains resulted in extremely high sediment supply into the north sea basin. this led to marked progradation from both the north and south. 100 km a 100 km b bulletin 22_ gsb191-indhold 04/03/11 12.41 side 86 acknowledgements this project would not have been possible without the enthusiastic participation of the geologists working in the municipalities of vejle, ringkøbing, ribe, århus and sønderjylland (subsequently restructured into mc ribe, mc ringkøbing, mc århus and region syddanmark). the fundamental sedimentological and stratigraphic research on which this study is based was supported financially by the carlsberg foundation. statens naturvidenskabelige forskningsråd (danish natural science research council) generously supported the drilling of the cored borehole at sdr. vium. keld rømer rasmussen kindly provided the gamma-ray log from the addit gravel-pit borehole. tibor czako and k. ingemann schnetler are thanked for fruitful discussions on miocene geology. claus heilmannclausen is thanked for guidance to the localities in the limfjorden area. the authors thank claus heilmannclausen and dan evans for thorough and constructive reviews; 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compressional deformation: the alpine foreland and other examples. tectonophysics 252, 7– 59. 92 bulletin 22_ gsb191-indhold 04/03/11 12.41 side 92 bulletin 22_ gsb191-indhold 04/03/11 12.41 side 93 de nationale geologiske undersøgelser for danmark og grønland (geus) geological survey of denmark and greenland øster voldgade 10, dk-1350 copenhagen k denmark the series geological survey of denmark and greenland bulletin started in 2003 and replaced the two former bulletin series of the survey, viz. geology of greenland survey bulletin and geology of denmark survey bulletin. some of the twenty-one volumes published since 1997 in those two series are listed on the facing page. the present series, together with geological survey of den mark and greenland map series, now form the peer-reviewed scientific series of the survey. geological survey of denmark and greenland bulletin 1 the jurassic of denmark and greenland, 948 pp. (28 articles), 2003. edited by j.r. ineson & f. surlyk. 500.00 2 fish otoliths from the paleocene of denmark, 94 pp., 2003. by w. schwarzhans. 100.00 3 late quaternary environmental changes recorded in the danish marine molluscan faunas, 268 pp., 2004. by k.s. pedersen. 200.00 4 review of survey activities 2003, 100 pp. (24 articles), 2004. edited by m. sønderholm & a.k. higgins. 180.00 5 the jurassic of north-east greenland, 112 pp. (7 articles), 2004. edited by l. stemmerik & s. stouge. 160.00 6 east greenland caledonides: stratigraphy, structure and geochronology, 93 pp. (6 articles), 2004. edited by a.k. higgins and f. kalsbeek. 160.00 7 review of survey activities 2004, 80 pp. (19 articles), 2005. edited by m. sønderholm & a.k. higgins. 180.00 8 structural analysis of the rubjerg knude glaciotectonic complex, vendsyssel, northern denmark, 192 pp., 2005. by s.a.s. pedersen. 300.00 9 scientific results from the deepened lopra-1 borehole, faroe islands, 156 pp. (11 articles), 2006. edited by j.a. chalmers & r. waagstein. 240.00 10 review of survey activities 2005, 68 pp. (15 articles), 2006. edited by m. sønderholm & a.k. higgins. 180.00 11 precambrian crustal evolution and cretaceous–palaeogene faulting in west greenland, 204 pp. (12 articles), 2006. edited by a.a. garde & f. kalsbeek. 240.00 12 lithostratigraphy of the palaeogene – lower neogene succession of the danish north sea, 77 pp., 2007. by p. schiøler, j. andsbjerg, o.r. clausen, g. dam, k. dybkjær, l. hamberg, c. heilmann-clausen, e.p. johannessen, l.e. kristensen, i. prince & j.a. rasmussen. 240.00 13 review of survey activities 2006, 76 pp. (17 articles), 2007. edited by m. sønderholm & a.k. higgins. 180.00 14 quaternary glaciation history and glaciology of jakobshavn isbræ and the disko bugt region, west greenland: a review, 78 pp., 2007. by a. weidick & o. bennike. 200.00 15 review of survey activities 2007, 96 pp. (22 articles), 2008. edited by o. bennike & a.k. higgins. 200.00 16 evaluation of the quality, thermal maturity and distribution of potential source rocks in the danish part of the norwegian–danish basin, 66 pp., 2008. by h.i. petersen, l.h. nielsen, j.a. bojesen-koefoed, a. mathiesen, l. kristensen & f. dalhoff. 200.00 17 review of survey activities 2008, 84 pp. (19 articles), 2009. edited by o. bennike, a.a. garde & w.s. watt. 200.00 18 greenland from archaean to quaternary. descriptive text to the 1995 geological map of greenland, 1:2 500 000. 2nd edition, 126 pp., 2009. by n. henriksen, a.k. higgins, f. kalsbeek & t.c.r. pulvertaft. 280.00 19 lithostratigraphy of the cretaceous–paleocene nuussuaq group, nuussuaq basin, west greenland, 171 pp., 2009. by g. dam, g.k. pedersen, m. sønderholm, h.h. midtgaard, l.m. larsen, h. nøhr-hansen & a.k. pedersen. 300.00 20 review of survey activities 2009, 106 pp. (23 articles), 2010. edited by o. bennike, a.a. garde & w.s. watt. 220.00 21 exploration history and place names of northern east greenland, 368 pp., 2010. by a.k. higgins. 200.00 22 lithostratigraphy of the upper oligocene – miocene succession of denmark, 92 pp., 2010. by e.s. rasmussen, k. dybkjær & s. piasecki. bulletin 22_ gsb191-indhold 04/03/11 12.41 side 94 geological survey of denmark and greenland map series 1 explanatory notes to the geological map of greenland, 1:500 000, humboldt gletscher, sheet 6, 48 pp., 2004. by p.r. dawes 280.00 2 explanatory notes to the geological map of greenland, 1:500 000, thule, sheet 5 (1991), 97 pp. + map, 2006. by p.r. dawes. 300.00 3 explanatory notes to the geological map of greenland, 1:100 000, ussuit 67 v.2 nord, 40 pp. + map, 2007. by j.a.m. van gool & m. marker. 280.00 4 descriptive text to the geological map of greenland, 1:500 000, dove bugt, sheet 10, 32 pp. + map, 2009. by n. henriksen & a.k. higgins 240.00 5 descriptive text to the geological map of greenland, 1:100 000, kangaatsiaq 68 v.1 syd and ikamiut 68 v.1 nord, 41 pp. + 2 maps, 2010. by a.a. garde & j.a. hollis. geology of greenland survey bulletin (discontinued) 179 the citronen fjord massive sulphide deposit, peary land, north greenland: discovery, stratigraphy, mineralization and structural setting, 40 pp., 1998. by f.w. van der stijl & g.z. mosher. 200.00 180 review of greenland activities 1997, 176 pp. (26 articles), 1998. edited by a.k. higgins & w.s. watt. 200.00 181 precambrian geology of the disko bugt region, west greenland, 179 pp. (15 articles), 1999. edited by f. kalsbeek. 240.00 182 vertebrate remains from upper silurian – lower devonian beds of hall land, north greenland, 80 pp., 1999. by h. blom. 120.00 183 review of greenland activities 1998, 81 pp. (10 articles), 1999. edited by a.k. higgins & w.s. watt. 200.00 184 collected research papers: palaeontology, geochronology, geochemistry, 62 pp. (6 articles), 1999. 150.00 185 greenland from archaean to quaternary. descriptive text to the geological map of greenland, 1:2 500 000, 93 pp., 2000. by n. henriksen, a.k. higgins, f. kalsbeek & t.c.r. pulvertaft. 225.00 186 review of greenland activities 1999, 105 pp. (13 articles), 2000. edited by p.r. dawes & a.k. higgins. 225.00 187 palynology and deposition in the wandel sea basin, eastern north greenland, 101 pp. (6 articles), 2000. edited by l. stemmerik. 160.00 188 the structure of the cretaceous–palaeogene sedimentary-volcanic area of svartenhuk halvø, central west greenland, 40 pp., 2000. by j. gutzon larsen & t.c.r. pulvertaft. 130.00 189 review of greenland activities 2000, 131 pp. (17 articles), 2001. edited by a.k. higgins & k. secher. 160.00 190 the ilímaussaq alkaline complex, south greenland: status of mineralogical research with new results, 167 pp. (19 articles), 2001. edited by h. sørensen. 160.00 191 review of greenland activities 2001, 161 pp. (20 articles), 2002. edited by a.k. higgins, k. secher & m. sønderholm. 200.00 geology of denmark survey bulletin (discontinued) 36 petroleum potential and depositional environments of middle jurassic coals and non-marine deposits, danish central graben, with special reference to the søgne basin, 78 pp., 1998. by h.i. petersen, j. andsbjerg, j.a. bojesen-koefoed, h.p. nytoft & p. rosenberg. 250.00 37 the selandian (paleocene) mollusc fauna from copenhagen, denmark: the poul harder 1920 collection, 85 pp., 2001. by k.i. schnetler. 150.00 prices are in danish kroner exclusive of local taxes, postage and handling note that information on the publications of the former geological survey of denmark and the former geological survey of greenland (amalgamated in 1995 to form the present geological survey of denmark and greenland) can be found on the survey’s website: www.geus.dk bulletin 22_ gsb191-indhold 04/03/11 12.41 side 95 bulletin 22_ gsb191-indhold 04/03/11 12.41 side 96 geological survey of denmark and greenland bulletin 6, 29-39 29 the neoproterozoic rivieradal group of kronprins christian land, eastern north greenland m. paul smith, a.k. higgins, n.j. soper and martin sønderholm the rivieradal group, formally defined here, is confined to the vandredalen thrust sheet of the caledonian orogen in kronprins christian land, eastern north greenland. it comprises a succession of neoproterozoic siliciclastic sediments that represent the fill of a half-graben basin. the syn-rift rivieradal group is overlain by post-rift sediments of the hagen fjord group. the latter succession is present in both the thrust sheet and the caledonian foreland to the west. in the foreland, where the rivieradal group is not represented, the hagen fjord group disconformably overlies palaeoproterozoic–mesoproterozoic sandstones of the independence fjord group. keywords: caledonian, north greenland, precambrian, proterozoic, stratigraphy m.p.s., lapworth museum, school of geography, earth and environmental sciences, university of birmingham, edgbaston, birmingham b15 2tt, uk. e-mail: m.p.smith@bham.ac.uk a.k.h. & m.s., geological survey of denmark and greenland, øster voldgade 10, dk-1350, copenhagen k, denmark. n.j.s., gams bank, threshfield, skipton bd23 5np, uk. also affiliated with: department of geology, university college, galway, ireland. kronprins christian land lies at the northern termination of the east greenland caledonides, and constitutes a key area for studies of the western border zone of the orogen (fig. 1). this region exposes continuous sections from the undisturbed foreland in the west, across parautochthonous foreland affected by folding and thin-skinned thrusting to allochthonous thrust sheets in the east (higgins et al. 2001a, b). the foreland comprises three principal lithostratigraphical divisions: (1) palaeoproterozoic–mesoproterozoic sandstones of the independence fjord group, the mesoproterozoic zig-zag dal basalt formation and associated dolerites (midsommersø dolerite formation); (2) neoproterozoic shallow marine sediments of the hagen fjord group (sønderholm & jepsen 1991); (3) cambrian–silurian shelf sediments of the franklinian basin (higgins et al. 1991). the hagen fjord group is also represented within the allochthon, in the vandredalen thrust sheet, where it overlies the clastic sediments of the rivieradal group, which are the subject of this paper. in kronprins christian land the parautochthonous lower palaeozoic sediments lie in the foot wall of the vandredalen thrust sheet, and are deformed by a series of thin-skinned thrusts that constitute a duplex below the vandredalen thrust (figs 1, 2). this parautochthonous area extends as a 30–50 km wide belt to the west of the 200 km long, n–s-trending vandredalen thrust front. the vandredalen thrust displays a classical staircase trajectory with very long (20+ km) flats developed in dolomitic horizons, and ramps developed in the more resistant subtidal carbonate units of the franklinian basin succession (cf. smith et al. 2004, this volume). the thrust roots to the east along the spærregletscher – hekla sund lineament and has a total westward displacement estimated at c. 40 km, of which c. 18 km are taken up in the thin-skinned parautochthonous belt (higgins et al. 2001b, 2004). the geological survey of denmark and greenland bulletin 6, 29–39 © geus, 2004 geus bulletin 6.pmd 10-02-2005, 09:5329 30 ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲▲ ▲ ▲ ▲ ro mer sø i n go lf fjord amdrup land holm land hovgaard ø kap bernhoft dijmphna sund hek la sund centrumsø sk a l l in g e n syd vej dal sp t sp t r iv ierada l 20°w 80°n 81°n blåsø sk jo ld un ge elv nioghalvfje rdsfjorden græ sel v sæ faxi elv d an m ar k fj or d va nd re d al en th ru st va nd re da le n kap holbæk p c m a hj h bs fl d v lambert land mfig. 3 wandel sea basin sequence (post-caledonian) samuelsen høj formation lauge koch land formation odins fjord formation turesø formation wandel valley formation kap holbæk formation crystalline basement thrust fault, shear zone børglum river and sjælland fjelde formations fyns sø, kap bernhard, campanuladal fms hagen fjord gp rivieradal group independence fjord gp and basaltic formations vandredalen thrust hagen fjord group ▲ ▲ ▲ ▲ vandredalen thrust sheet ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲▲ ▲ ▲ c a le d o n id es greenland i 25 km geus bulletin 6.pmd 10-02-2005, 09:5330 31 vandredalen thrust sheet, in turn, is structurally overlain by a thrust sheet that transported palaeoproterozoic to mesoproterozoic clastic and volcanic rocks westwards. the allochthonous quartzites have traditionally been viewed as equivalents of the independence fjord group on the foreland (see figs 1, 2), but shrimp isotopic studies on rhyolites interbedded with the quartzites that yielded an age of 1740 ma (kalsbeek et al. 1999) have cast some doubt on this interpretation (see also pedersen et al. 2002). still farther to the east, higher thick-skinned thrust sheets incorporate crystalline basement gneisses. this paper provides a formal stratigraphic basis, at group level, for the rocks that underlie the hagen fjord group within the vandredalen thrust sheet, and collates available field data regarding this succession. as noted below, formal definition of the constituent formations awaits more detailed field analysis of the rivieradal group. the informal units of fränkl (1954, 1955) were found to be generally usable as field divisions, although there is uncertainty in places about their correlation. stratigraphy of the vandredalen thrust sheet restoration of the displacement associated with the vandredalen thrust sheet demonstrates that much of the sediment within the thrust sheet was deposited in an east-facing extensional half-graben (hekla sund basin) that originally lay immediately to the east of the spærregletscher – hekla sund lineament (higgins et al. 2001b). this basin fill was thrust out of the halfgraben and transported westwards on the vandredalen thrust (fig. 2). these syn-rift sediments are assigned to the rivieradal group, which is formally erected below. they are overlain by sediments of the hagenfjord group, which were deposited during postrift thermal subsidence and extended westwards beyond the confines of the hekla sund basin. the hagen fjord group, in consequence, is present both in the foreland and the hanging wall of the vandredalen thrust whereas the rivieradal group is restricted to the hanging wall (higgins et al. 2001b). the hagen fjord group within the vandredalen thrust sheet is represented by the campanuladal, kap bernhard and fyns sø formations. the campanuladal formation comprises 200 m of variegated sandstones, siltstones and mudstones of generally similar appearance to successions of the same unit in the foreland around danmark fjord.a lower, greenish weathering unit containing parallel and trough cross-laminated sandstones, and mudstones with desiccation cracks, is overlain by an upper, dark red weathering unit dominated by calcareous mudstones with some trough cross-bedded sandstones (jepsen&sønderholm1994). thekapbernhard formation comprises brownish red weathering, finely laminated algal limestones and the fyns sø formation contains pale stromatolitic dolostones; both correspond very closely to their counterparts in the foreland. in the northernmost part of the outcrop area of the vandredalen thrust sheet, in finderup land (fig. 1), the fyns sø formation is unconformably overlain by sandstones of the kap holbæk formation. this latter unit was included within the redefined hagen fjord group of clemmensen & jepsen (1992), but stratigraphic data obtained during the 1994–1995 field seasons demonstrated that a significant hiatus occurs between the two units. in particular, a well-developed palaeokarst horizon developed at the top of the fyns sø formation is infilled by the kap holbæk formation (smith et al. 1999). it is probable that the carbonates of the fyns sø formation are correlatives of similar late riphean to sturtian units, which are widely developed in the north atlantic region, while the presence of deep skolithos burrows in the kap holbæk formation indicates a lower cambrian age (smith et al. 2004, this volume). since the hiatus between the two units probably spans the entire vendian, smith et al. (2004) proposed that the kap holbæk formation be removed from the hagen fjord group. facing page: fig. 1. geological map of kronprins christian land, eastern north greenland, and lambert land, north-east greenland. black oval symbols in the rivieradal group, in legend and on map, indicate conglomerates. bs, brede spærregletscher; d, ‘dunkeldal’; fl, finderup land; h, harefjeld; hj, hjørnegletscher; m, marmorvigen; pcma, prinsesse caroline mathilde alper; spt, spærregletscher thrust; v, vardedalen. on index map: i, independence fjord. see figs 2 and 3 for cross-sections along rivieradal. modified from rasmussen & smith (2001). note that the traditional interpretation of the allochthonous quartzites as equivalents of the foreland independence fjord group (as depicted here and on fig. 2), has been brought into doubt by the 1740 ma age on interbedded rhyolites (see discussion in kalsbeek et al. 1999 and pedersen et al. 2002). geus bulletin 6.pmd 10-02-2005, 09:5331 32 rivieradal group new group history. the proterozoic rocks of kronprins christian land were first examined systematically by geologists of lauge koch’s 1926–1958 expeditions. the region around danmark fjord was documented by adams & cowie (1953), and that around centrumsø by fränkl (1954, 1955). fränkl recognised that the neoproterozoic succession could be divided into autochthonous and allochthonous parts, the two separated by a major thrust upon which his ‘main nappe’ was transported. the metasediments of the nappe were divided into a lower, more metamorphosed part, comprising the stenørkenen phyllites (> 1000 m) and the sydvejdal marbles (100–400 m), and an upper less metamorphic part. the latter included, from base to top, the taagefjeldene greywackes (> 700 m) with a layer of alum shales at the base (c. 150 m), the rivieradal sandstones (1000–2000 m), the ulvebjerg sandstones & tillites (20–35 m), the red, shaly campanuladal limestone and the fyns sø formation. the two last-named units were also identified as lying in the foot wall of the nappe (fränkl 1954, 1955). the succession present in the ‘main nappe’ was recognised as having been deposited in a basin located to the east, which fränkl (1955) termed the ‘hekla sund basin’; the latter term is retained here for the depositional basin in which the rivieradal group accumulated. fig. 2. schematic nw–se cross-sections of the vandredalen rift system, approximately along the section line of fig. 3 shown on fig. 1, but with extensions to both south-east and north-west. a: present-day section, with overburden calculated from conodont alteration temperatures (after rasmussen & smith 2001). b: with displacements on the thrusts restored, illustrating the fault-bounded control of the west margin of the hekla sund basin. modified from higgins et al. (2001b). spærregletscher thrust if rg 22 km 18 km spt 50 ? km 100 ? km st hf if o–s o–s o–s hfvt vt o–s upper allochthon– higher thrust sheets comprising ordovician–silurian, hagen fjord group, rivieradal group (distal part of hekla sund basin), independence fjord group and associated volcanic rocks and crystalline basement. estimated overburden st vt spt spærregletscher thrust vandredalen thrust sole thrust rg if cb hf o–s ordovician–silurian hagen fjord group rivieradal group crystalline basement independence fjord group and associated volcanic rocks 5 km 0 10 20 km hf present day rg cb cb if if if if cb cb hekla sund basin rivieradal group succession) a b wnw ese conglomerate hekla sund basin (rivieradal group) sandstone mudstone and calcareous mudstone geus bulletin 6.pmd 10-02-2005, 09:5432 33 haller (1961, 1971) erected the hagen fjord group for proterozoic sediments of the autochthon and allochthon in kronprins christian land, including volcanic rocks now referred to the zig-zag dal basalt formation, hekla sund formation and aage berthelsen gletscher formation, and lower cambrian carbonates now referred to the portfjeld formation. clemmensen & jepsen (1992) restricted the group, following work by the geological survey of greenland in 1978–1980, to include only the neoproterozoic shallow water carbonate and siliciclastic succession present in the area between lambert land and independence fjord. additional formations were also erected by clemmensen & jepsen, to improve the stratigraphical framework of the group. hurst & mckerrow (1981a, b), on the basis of reconnaissance field work in kronprins christian land in 1980 by the former geological survey of greenland (ggu), concluded that all of the units beneath the campanuladal formation within the main thrust sheet were representative of a single unit of deepwater turbidites, mud and resedimented conglomerates; they collectively referred to this succession as the ‘rivieradal sandstones’. although sedimentological interpretations of the succession have been refined (see below), this is the concept that we here formalise as the rivieradal group. hurst & mckerrow (1981a, b) and hurst et al. (1985) recognised a number of thrust sheets in the kronprins christian land sector of the east greenland caledonides. these included the vandredalen nappe (which corresponds to fränkl’s ‘main nappe’), the finderup land nappe and the sæfaxi elv nappe. the finderup land nappe was a geographically isolated structure rimming the western edge of the main ice sheet, and a succession that incorporated the campanuladal, fyns sø and kap holbæk formations was recognised. subsequent regional mapping by ggu in 1994–1995 has demonstrated that the finderup land nappe is a northward continuation of the vandredalen nappe (now the vandredalen thrust sheet) in which only the higher stratigraphic levels are preserved. the sæfaxi elv nappe was considered by hurst & mckerrow to be a thrust sheet containing allochthonous early palaeozoic sediments that were deep-water equivalents of the platform succession, and these were assigned to the harefjeld formation by hurst (1984). however, rasmussen & smith (1996) demonstrated that these sediments were highly strained equivalents of the platform succession, and that the lower contact with the underlying fyns sø formation was an unconformity and not a thrust. the ‘sæfaxi elv nappe’ is thus a succession of parautochthonous foreland carbonates deformed in the foot wall of the vandredalen thrust, and the concept of an independent thrust sheet (and of the harefjeld formation) has been abandoned (rasmussen & smith 1996). name. the group takes its name from rivieradal, the e–w-trending valley south-east of centrumsø that contains the most complete section through the unit (higgins & soper 1994, 1995). type area and reference sections. the type area for the group is rivieradal itself, where an excellent reference section through the upper part of the group occurs and the lower part, although highly deformed, is also present. additional reference sections through parts of the group are available through ‘dunkeldal’ (between northern vandredalen and inner ingolf fjord; fig. 1), and along much of the western side of vandredalen between the western end of ingolf fjord and romer sø. thickness. hurst et al. (1985) estimated a thickness of 2.5 km for the ‘rivieradal sandstones’. structural studies in rivieradal and ‘dunkeldal’, carried out during the 1994 and 1995 field seasons, suggest that the combined thickness of the highly deformed lower part of the rivieradal group and the less deformed upper part is substantially higher, in the order of 7.5–10 km (higgins et al. 2001b). lithology, facies associations and depositional environments. the sedimentological and stratigraphical variations within the rivieradal group should be viewed in the context of the internal structure of the vandredalen thrust sheet, which is best exposed in the 35 km long section along rivieradal between the thrust sheet front at the east end of centrumsø and its trailing edge near marmorvigen (fig. 3). the western frontal region of the thrust sheet is characterised by simple, large-scale folds developed in the fyns sø and kap bernhard formations; these overlie the rivieradal group sediments, which are well exposed throughout the valley of rivieradal. steep to vertical dips characterise much of the western third of the section, with occasional west-facing fold pairs showing flat common limbs. the central third of the section begins with an abrupt change to tightly developed folds, which become progressively more intense eastwards with the inclination of the axial surfaces decreasing geus bulletin 6.pmd 10-02-2005, 09:5433 34 from moderate eastward dips to almost horizontal. in the eastern third of the section deformation is intense with complete erasure of sedimentary way-up indicators in long-limbed isoclinal folds; the units exposed here correspond to the most distal parts of the rivieradal group. the fold style and orientation of strata in the rivieradal section suggest that the vandredalen thrust follows a series of ramps and flats, and can be viewed as sampling an oblique segment of the half-graben hekla sund basin in which the rivieradal group accumulated. the oldest and most distal parts of the succession are therefore preserved in the eastern part of the thrust sheet, and the younger and more proximal lithofacies are present in the west. the present-day distribution of remnants of the rivieradal group succession indicates that the original hekla sund basin must have been at least 200 km long from north to south and 50 km wide from east to west. the rivieradal group is lithologically variable and possesses a strong proximal to distal polarity. one of the most distinctive lithologies present is a coarse conglomerate, which occurs repeatedly along the leading edge of the vandredalen thrust sheet. substantial conglomerate units are present in three discrete areas (fig. 1). the northernmost is along a 15–20 km strike section on the west side of romer sø, the second is a 20 km long strike section on the west side of central vandredalen, and the southernmost is found in southern skallingen near blåsø. the best known of the conglomeratic successions is that in central vandredalen where several thick sections have been measured. quartzite clasts (90–95%) dominate everywhere over dolerite clasts (5–10%), and are probably derived from the independence fjord group and the midsommersø dolerite formation; these units are presumed to have been exposed to active erosion to the west of the basin. clasts vary in size from a few decimetres to well over a metre in the thicker beds, with occasional outsize clasts as much as 3–4 m across; these large clasts are indisputably proximal. viewed from a distance, the thick conglomerate units appear to have a lensoid form, suggesting deposition in a series of nested channels. in e–w valley sections to the west of romer sø, rapid lateral facies changes are well exposed. thick conglomerate-dominated units pass eastwards, over asa nd st on e m ud st on e m ud st on e, s ilt st on e an d sa nd st on e c ar bo na te a nd ca lc ar eo us m ud st on e lo w er p al ae oz oi c pl at fo rm t u , t ur es ø fo rm at io n; b r , b ør gl um r iv er f or m at io n; s f, sj æ lla nd f je ld e fo rm at io n; w v, w an de l v al le y fo rm at io n r iv ie ra da l g ro up ( pr ot er oz oi c) o th er p ro te ro zo ic a nd l ow er p al ae oz oi c un its : fs , f yn s sø f or m at io n; k b, k ap b er nh ar d fo rm at io n; c d , c am pa nu la da l f or m at io n; if , i nd ep en de nc e fj or d g ro up 0 5 10 k m 1 km if if k b fs sf fs w v k b br pa ra ut oc ht ho no us th ru st b el t va nd re da le n th ru st fr on t h ag en f jo rd g ro up r iv ie ra da l g ro up m ar m or vi ge n es e w n w br c d t u w v ? v a n d r ed a le n t h r u st fig. 3. cross-section through the vandredalen thrust sheet along rivieradal. see fig. 1 for section line. modified from higgins et al. (2001b). geus bulletin 6.pmd 10-02-2005, 09:5434 35 distance of 1–2 km, into upward-thickening packets of sandstone in which the individual beds coarsen upwards, sometimes into conglomerate. these in turn pass farther eastwards, over a similar distance, into upward-thickening and coarsening mudstone-siltstonesandstone packets. in the southern area of conglomerates near blåsø, rounded granite and quartz pebbles (up to 20%) make an appearance, although most clasts are again quartzite and dolerite; a deeper erosion level was evidently reached in the southern source region. the geometry of the conglomerate deposits, together with their discrete occurrences, suggests the presence of three discrete fan delta systems that acted as major feeder distributary systems on the western side of the original basin. while the three main conglomerate developments are all in the upper part of the rivieradal group succession, they may not be at exactly the same stratigraphical level. input might have been via a single major fluvial system, which varied in position with time, and if so the blåsø fan delta with its crystalline clasts may be the youngest of these. the repeated cycles of conglomeratic deposition, and the upward-coarsening sandstone cycles in other areas at the frontal part of the thrust sheet, may have been controlled by displacements on the basin-margin fault system. in ‘dunkeldal’, a valley on the east side of vandredalen, a total thickness of 3000 m has been measured in a continuously exposed section (lower part of measured section in fig. 4). the basal 200 m of this section lie above a thrust contact with ordovician carbonates, and comprise strongly sheared conglomerates. the conglomerates are overlain by a 500 m thick phyllite-dominated unit (‘stenørkenen phyllites’ of fränkl 1955), and then by over 2200 m of sandstone turbidites interbedded with dark pyritic mudstones (‘taagefjeldene greywackes’ of fränkl). laterally and vertically, this sandstone-dominated succession grades into homogeneous black mudstones, and equivalent phyllitic rocks, which are widely exposed around the innermost branch of ingolf fjord. the lack of marker horizons and non-exposure in the flat valley bottom of vandredalen makes thickness estimates in these areas difficult; thus the notional gap of 350 m in the measured section of fig. 4 may in fact correspond to several kilometres of section. on the west side of vandredalen, to the north-west of innermost ingolf fjord, a 900 m thick succession comprises coarsening-upwards sequences of parallel laminated mudstone, lenticular and wavy-bedded mudstones, sandstone-dominated heterolithic sediments with parallel lamination, trough cross-lamination and hummocky cross-stratification, overlain by trough and planar cross-bedded sandstones with occasional herringbone cross-bedding (upper part of section in fig. 4). these correspond to the original ‘rivieradal sandstones’ of fränkl (1955). current directions are predominantly towards the north-east and are interpreted as the products of stormand tidedominated shallow marine deposition. this facies association in the upper part of the rivieradal group can be recognised throughout the outcrop area, from romer sø in the north to blåsø in the south. on the west side of vandredalen, the succession of fig. 4 continues with about 300 m of interfingering conglomerates and sandstones (the ‘ulvebjerg sandstones and tillites’ of fränkl – equivalent to the conglomeratic developments described above), which are overlain by sediments referred to the hagen fjord group. the latter shallow marine succession constitutes the postrift fill of the basin. the most distal representatives of the rivieradal group are seen in the valley of rivieradal itself. at the eastern end of the valley, pelitic slates with sandstones form the coastal mountains south of the mouth of rivieradal. these are overlain farther to the west by pelitic and calcareous slates and siltstones with prominent yellow-weathering carbonate units, which correspond to the ‘sydvejdal marbles with chloritic shales’ of fränkl (1955). this unit is overlain, in turn, by phyllites and turbidites corresponding to those seen in the ‘dunkeldal’ section (fig. 4). overall, the rivieradal group is characterised by point sources of sediment input which generated substantial conglomerate fan deltas, and which are associated with sandy, proximal turbidites. between the fans and in the eastern (distal) part of the basin, sedimentation was dominated by mud and calcareous mud. as the basin filled, the depositional style switched from deep to shallow marine, and less localised, more laterally persistent, tidal and storm-dominated deposition began to predominate. boundaries. since the group is restricted to the vandredalen thrust sheet, the rivieradal group is everywhere bounded on its lower surface by the vandredalen thrust; a stratigraphic base to the group has not been identified within the thrust sheet. the upper boundary of the group is placed where sandstones and laterally equivalent conglomerates are overlain by a characteristic 200 m variegated unit comprising a geus bulletin 6.pmd 10-02-2005, 09:5435 36 1200 2500 3700 3600 3500 3400 3300 3200 3100 3000 2900 2800 2700 2600 2500 3800 3900 4000 4100 4200 4300 4400 4500 4600 4700 4800 4900 5000 2400 2300 2200 2100 2000 1900 1800 1700 1600 1500 1400 1300 1100 1000 900 800 700 600 500 400 300 200 1000 ? ? 'taagefjeldene greywackes' massive ta,b sandstone turbidites and black pyritic mudstones 'stenørkenen phyllites' ? black phyllites, pyritic 'taagefjeldene greywackes' not exposed black mudstones with thin sandstones sandstone dominated coarsening-upward sequences fyns sø fm pale dolostone kap bernhard fm red limestone campanuladal fm. variegated sandstones and mudstones 'ulvebjerg sandstones' laterally interfingering conglomerates and sandstones 'rivieradal sandstones' mudstone dominated coarsening-upward sequences basal conglomerate basal thrust m ud cracks c rossbedding parallel lam ination 'rivieradal sandstones' 'taagefjeldene greywackes' g e u s b ulletin 6.pm d 10-02-2005, 09:54 36 37 lower greenish sandstone and mudstone interval overlain by a dark red calcareous mudstone-dominated interval. the greenish sandstone and mudstone interval contains parallel and trough cross-laminated sandstones and abundant desiccation cracks. the variegated unit is identified as the campanuladal formation of the foreland and, as in the foreland, is overlain by brownish red weathering microbially laminated limestones of the kap bernhard formation (fig. 4). distribution. the rivieradal group is restricted to the vandredalen thrust sheet, and crops out in a broad zone extending from romer sø southwards along vandredalen; a further broad zone extends from the eastern end of centrumsø southwards through rivieradal and skallingen to blåsø (fig. 1). a narrow strip of outcrops, in the hanging wall of the main thrust ramp, extends from vardedalen (on the north side of central ingolf fjord) southwards to brede spærregletscher and along the west side of hekla sund to marmorvigen. the southernmost outcrops of the rivieradal group are present in nunataks at the westernmost extremity of lambert land (fig. 1). geological age. the group is older than the hagen fjord group, specifically the kap bernhard and fyns sø formations, thought to be of probable riphean age (smith et al. 1999). however, frederiksen (2000) has suggested the hagen fjord group is sturtian, and equivalent to the andrée land group of the eleonore bay supergroup. the rivieradal group post-dates the independence fjord group, midsommersø dolerite formation and zig-zag dal basalt formation, all of which are represented by clasts within the conglomeratic units. the dolerites of the midsommersø dolerite formation were originally dated at c. 1230 ma by kalsbeek & jepsen (1984), but a recent baddeleyite age on a dolerite of 1380 ma has been obtained by upton et al. (in press); this provides a maximum age limit for deposition. the rivieradal group was thus deposited in the interval between 1380 ma and ~ 700 ma. g. vidal (in hurst et al. 1985) recorded ‘several comparatively well-preserved specimens’ of acritarchs from the upper part of the rivieradal group which were thought indicative of an upper proterozoic age. in particular, a single specimen of chuaria circularis was considered to be indicative of an upper riphean age since, elsewhere in scandinavia and north america, the species occurs at around 800 ma. taken together with the evidence for a conformable upper boundary, this suggests that the rivieradal group was deposited in the younger part of the broad age range outlined above. subdivision. fränkl (1954, 1955) recognised five units within the succession now assigned to the rivieradal group: ‘stenørkenen phyllites’, ‘sydvejdal marbles with chloritic shales’, ‘taagefjeldene greywackes’ with a layer of alum shales at the base, ‘rivieradal sandstones’, and the ‘ulvebjerg sandstones and tillites’. these units were used as field terms during the 1994–1995 field seasons and proved to be recognisable throughout the area, although this is at least partly because they represent lithological types rather than coherent and homologous stratigraphic units. the ‘stenørkenen phyllites’ are present in rivieradal and similar phyllites are present throughout the region, although thickness estimates and correlation is hampered by the paucity of marker horizons. fränkl (1955) estimated a thickness of > 1000 m. the ‘sydvejdal marbles with chloritic shales’ are also present in eastern rivieradal, and include pelitic and calcareous slates and siltstones with prominent yellow-weathering carbonate units. fränkl (1955) estimated a thickness of 100–400 m. the ‘taagefjeldene greywackes’ are best seen in the section through ‘dunkeldal’ (fig. 4), where jepsen & sønderholm (1994) recorded a thickness of over 2200 m. this compares with fränkl’s (1954, 1955) estimate of > 700 m, of which 150–200 m were alum shales; multicoloured friable shales cover large areas around the head of ingolf fjord, and pass laterally (northwards) along vandredalen into a sandstonesiltstone-mudstone succession. the ‘rivieradal sandstones’ sensu fränkl have a thickness of 1400 m, measured in a section across skallingen for which no base was present (leslie & jepsen 1995), and jepsen & sønderholm (1994) measured a thickness of 900 m on the western side of vandredalen (fig. 4). fränkl (1954) estimated a range of 1000–2000 m, which appears to be the right order of magnitude. the ‘tillites’ of fränkl’s (1954, 1955) ‘ulvebjerg sandstones and tillites’ unit were relatively soon re-interpreted as non-glacial facing page: fig. 4. simplified measured section through part of the synrift rivieradal group succession, the lower part in ‘dunkeldal’, and the upper part from the west side of vandredalen. the corresponding lithostratigraphical terms of fränkl (1955) are indicated. the designation ‘t a,b ’ refers to bouma cycle intervals in the sandstone turbidites. the campanuladal, kap bernhard and fyns sø formations form part of the post-rift hagen fjord group. slightly modified from jepsen & sønderholm (1994). geus bulletin 6.pmd 10-02-2005, 09:5437 38 conglomerate horizons (haller 1971; hurst & mckerrow 1981a), and they are one of the most spectacular stratigraphic developments within the group. on the western side of vandredalen, a sandstonedominated unit at least 300 m thick can be seen to pass laterally into thick conglomerates up to 500 m thick (jepsen & sønderholm 1994). as noted above, three of these major conglomerate developments are present within the region and represent fan deltas that supplied sediment to the basin. although these observations indicate that fränkl’s units may in time form the basis for a stratigraphic framework, any formal definition must await more detailed investigation of the rivieradal group. acknowledgements we are grateful to the referees, l. clemmensen and j.s. peel, for their helpful comments. references adams, p.j. & cowie, j.f. 1953: a geological reconnaissance of the region round the inner part of danmarks fjord, northeast greenland. meddelelser om grønland 111(7), 24 pp. clemmensen, l.b. & jepsen, h.f. 1992: lithostratigraphy and geological setting of upper proterozoic shelf deposits, hagen fjord group, eastern north greenland. rapport grønlands geologiske undersøgelse 157, 27 pp. fränkl, e. 1954: vorläufige mitteilung über die geologie von kronprins christian land (ne-grönland). meddelelser om grønland 116(2), 85 pp. fränkl, e. 1955: weitere beiträge zur geologie von kronprins christian land (ne-grönland, zwischen 80° und 80°30′n). meddelelser om grønland 103(7), 35 pp. frederiksen, k.s. 2000: a neoproterozoic carbonate ramp and base-of-slope succession, the andrée land group, eleonore bay supergroup, north-east greenland: sedimentary facies, stratigraphy and basin evolution, 242 pp. unpublished ph.d. thesis, university of copenhagen, denmark. haller, j. 1961: the carolinides: an orogenic belt of upper precambrian age in northeast greenland. in: raasch, g.o. (ed.): geology of the arctic 1, 153–159. toronto: toronto university press. haller, j. 1971: geology of the east greenland caledonides, 413 pp. london: interscience. higgins, a.k. & soper, n.j. 1994: the caledonian thrust belt of kronprins christian land. in: henriksen, n. (ed.): express report: eastern north greenland and north-east greenland 1994, 57–67. unpublished report, geological survey of greenland, copenhagen. higgins, a.k. & soper, n.j. 1995: tectonic and other observations in vandredalen and adjacent areas of southern kronprins christian land. in: higgins, a.k. (ed.): express report: eastern north greenland and north-east greenland 1995, 105– 114. unpublished report, geological survey of greenland, copenhagen. higgins, a.k., ineson, j.r., peel, j.s., surlyk, f. & sønderholm, m. 1991: lower palaeozoic franklinian basin of north greenland. in: peel, j.s. & sønderholm, m. (eds): sedimentary basins of north greenland. bulletin grønlands geologiske undersøgelse 160, 71–139. higgins, a.k., leslie, a.g. & smith, m.p. 2001a: neoproterozoic – lower palaeozoic stratigraphical relationships in the marginal thin-skinned thrust belt of the east greenland caledonides: comparisons with the foreland in scotland. geological magazine 138(2), 143–160. higgins, a.k., smith, m.p., soper, n.j., leslie, a.g., rasmussen, j.a. & sønderholm, m. 2001b: the neoproterozoic hekla sund basin, eastern north greenland: a pre-iapetan extensional sequence thrust across its rift shoulders during the caledonian orogeny. journal of the geological society (london) 158, 487–499. higgins, a.k., soper, n.j., smith, m.p. & rasmussen, j.a. 2004: the caledonian thin-skinned thrust belt of kronprins christian land, eastern north greenland. in: higgins, a.k. & kalsbeek, f. (eds): east greenland caledonides: stratigraphy, structure and geochronology. geological survey of denmark and greenland bulletin 6, 41–56 (this volume). hurst, j.m. 1984: upper ordovician and silurian carbonate shelf stratigraphy, facies and evolution, eastern north greenland. bulletin grønlands geologiske undersøgelse 148, 73 pp. hurst, j.m. & mckerrow, w.s. 1981a: the caledonian nappes of eastern north greenland. nature 290, 772–774. hurst, j.m. & mckerrow, w.s. 1981b: the caledonian nappes of kronprins christian land, eastern north greenland. rapport grønlands geologiske undersøgelse 106, 15–19. hurst, j.m., jepsen, h.f., kalsbeek, f., mckerrow, w.s. & peel, j.s. 1985: origin of the caledonian nappes of eastern north greenland. in: gee, d.g. & sturt, b.a. (eds): the caledonide orogen: scandinavia and related areas, 1047–1063. london: john wiley. jepsen, h.f. & sønderholm, m. 1994: sedimentological studies of the hagen fjord group and ‘rivieradal sandstones’ (late proterozoic), eastern north greenland. in: henriksen, n. (ed.): express report: eastern north greenland and north-east greenland 1994, 39–48. unpublished report, geological survey of greenland, copenhagen. kalsbeek, f. & jepsen, h.f. 1984: the late proterozoic zig-zag dal basalt formation of eastern north greenland. journal of petrology 25, 644–664. kalsbeek, f., nutman, a.p., escher, j.c., friderichsen, j.d., hull, j.m., jones, k.a. & pedersen, s.a.s. 1999: geochronology of granitic and supracrustal rocks from the northern part of the east greenland caledonides: ion microprobe u-pb zircon ages. geology of greenland survey bulletin 184, 31–48. leslie, a.g. & jepsen, h.f. 1995: tectono-stratigraphic studies of the allochthonous ‘rivieradal sandstones’ and hagen fjord group (late proterozoic), kronprins christian land, eastern geus bulletin 6.pmd 10-02-2005, 09:5438 39 north greenland. in: higgins, a.k. (ed.): express report: eastern north greenland and north-east greenland 1995, 93– 103. unpublished report, geological survey of greenland, copenhagen. pedersen, s.a.s., craig, l.e., upton, b.g.j., rämö, o.t., jepsen, h.f. & kalsbeek, f. 2002: palaeoproterozoic (1740 ma) riftrelated volcanism in the hekla sund region, eastern north greenland: field occurrence, geochemistry and tectonic setting. precambrian research 114, 327–346. rasmussen, j.a. & smith, m.p. 1996: lower palaeozoic carbonates in eastern north greenland, and the demise of the ‘sæfaxi elv nappe’. bulletin grønlands geologiske undersøgelse 172, 49–54. rasmussen, j.a. & smith, m.p. 2001: conodont geothermometry and tectonic overburden in the northernmost east greenland caledonides. geological magazine 138(6), 687–698. smith, m.p., soper, n.j., higgins, a.k., rasmussen, j.a. & craig, l.e. 1999: palaeokarst systems in the neoproterozoic of eastern north greenland in relation to extensional tectonics on the laurentian margin. journal of the geological society (london) 156, 113–124. smith, m.p., rasmussen, j.a., robertson, s., higgins, a.k. & leslie, a.g. 2004: lower palaeozoic stratigraphy of the east greenland caledonides. in: higgins, a.k. & kalsbeek, f. (eds): east greenland caledonides: stratigraphy, structure and geochronology. geological survey of denmark and greenland bulletin 6, 5–28 (this volume). sønderholm, m. & jepsen, h.f. 1991: proterozoic basins of north greenland. in: peel, j.s. & sønderholm, m. (eds): sedimentary basins of north greenland. bulletin grønlands geologiske undersøgelse 160, 49–69. upton, b.g.j., rämö, o.t., heaman, l.m., blichert-toft, j., barry, t.l., kalsbeek, f. & jepsen, h.f. in press: the zig-zag dal basalts and associated intrusions of eastern north greenland: progressive mantle plume – lithosphere interaction. contributions to mineralogy and petrology. geus bulletin 6.pmd 10-02-2005, 09:5439 untitled 1 geological survey of denmark and greenland bulletin 8 · 2005 structural analysis of the rubjerg knude glaciotectonic complex, vendsyssel, northern denmark stig a. schack pedersen geological survey of denmark and greenland danish ministry of the environment 2 geological survey of denmark and greenland bulletin 8 keywords northern jylland, denmark, weichselian, glacial geology, glaciotectonics, thin-skinned thrust faulting, balanced cross-section, thrustfault dynamics, imbricate duplexes, mud diapirs, piggyback basins. cover the coastal clif f (99 m high at its highest point) at rubjerg knude on the west coast of vendsyssel, northern denmark. the lower two-thirds of the cliff, beneath the prominent dark sub-horizontal surface, forms part of the cross-section through the rubjerg knude glaciotectonic complex displaying imbricated thrust sheets composed of the lønstrup klint formation (bluish-grey colour) and the overlying rubjerg knude formation (yellow colour), both of late weichselian age. the thrust sheets are truncated by a glaciotectonic unconformity (the prominent surface), upon which the kattegat till formation is only preserved as a boulder bed due to subsequent aeolian erosion of the till matrix. the upper third of the clif f comprises recent aeolian dune sands that have accreted over the last 100 years and now encroach on the rubjerg knude lighthouse, the top of which is just visible above the clif ftop. photo: stig a. schack pedersen (august 1984). chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, geological institute, university of aarhus; minik rosing, geological museum, university of copenhagen; finn surlyk, geological institute, university of copenhagen scientific editor of this volume: jon r. ineson editorial secretaries: esben w. glendal and birgit eriksen illustrations: benny m. schark and alice rosenstand digital photographic work: benny m. schark graphic production: knud gr@phic consult, odense, denmark printers: schultz grafisk, albertslund, denmark manuscript submitted: 8 august 2003 final version approved: 11 february 2005 printed: 15 december 2005 this monograph has been accepted by the faculty of natural sciences, university of copenhagen, for public defence of the degree of doctor of science. issn 1604-8156 isbn 87-7871-168-1 geological survey of denmark and greenland bulletin the series geological survey of denmark and greenland bulletin replaces geology of denmark survey bulletin and geology of greenland survey bulletin. citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 8, 192 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk or geografforlaget aps rugårdsvej 55, dk-5000 odense c, denmark phone: +45 63 44 16 83, fax: +45 63 44 16 97, e-mail: go@geografforlaget.dk © danmarks og grønlands geologiske undersøgelse (geus), 2005 3 contents abstract . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9 introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 11 history of the present investigation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 11 objectives . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 14 glacial tectonics – concepts and models . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 14 previous conceptual models . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 14 thin-skinned thrust faulting: the concept . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16 thrust-fault modelling . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16 test model 1. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19 test model 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19 test model 3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 21 test model 4 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 21 test models: concluding remarks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 22 concept of balanced cross-section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 22 location and construction of the rubjerg knude cross-section . . . . . . . . . . . . . . . . . 23 location. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23 photogrammetric work. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23 digital editing . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25 construction of the balanced cross-section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 26 geological setting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 27 lithostratigraphy. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 33 skærumhede group . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 33 stortorn formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 35 lønstrup klint formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 39 upper weichselian lithostratigraphic units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 43 rubjerg knude formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 43 kattegat till formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 48 ribjerg formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 50 mid danish till formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 53 vendsyssel formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 55 structural description of sections . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 60 ulstrup section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 60 tectonic architecture . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 61 sedimentary units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 62 lønstrup klint formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 62 rubjerg knude formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 63 structures and breccias . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 63 thrust-zone breccias . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 64 foreland-dipping hanging-wall flat faults . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 64 collapse structure in the ulstrup rende . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 66 interpretation of structural development . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 66 stensnæs section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 70 tectonic architecture . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 70 sedimentary units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 72 lønstrup klint formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 72 rubjerg knude formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 73 structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 73 4 imbricate duplex folding . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 73 extensional faults . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 73 interpretation of structural development . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 78 martørv bakker section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 78 tectonic architecture . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 79 sedimentary units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 80 lønstrup klint formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 81 rubjerg knude formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 81 structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 82 imbricate duplexes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 82 normal fault . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 82 hydrodynamic brecciation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 82 interpretation of structural development . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 82 kramrende section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 84 tectonic architecture . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 85 sedimentary units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 86 lønstrup klint formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 87 rubjerg knude formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 87 structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 87 thrust faults . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 87 kramrende diapir . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 88 reverse faults . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 88 interpretation of structural development . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 89 brede rende section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 90 tectonic architecture . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 90 sedimentary units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 93 lønstrup klint formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 93 rubjerg knude formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 94 structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 94 diapir structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 96 brede rende normal fault . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 97 frost wedges . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 97 interpretation of structural development . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 97 sandrende section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 99 tectonic architecture . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 99 sedimentary units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 101 lønstrup klint formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 101 rubjerg knude formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 101 structures and breccias . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 103 normal faults . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 103 diapir structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 104 frost wedge . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 104 interpretation of structural development . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 104 stenstue rende section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 105 tectonic architecture . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 106 sedimentary units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 107 lønstrup klint formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 107 rubjerg knude formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 108 structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 108 thrust-fault structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 108 hanging-wall anticlines . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 109 5 normal faults . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 110 slump folding . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 110 interpretation of structural development . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 110 grønne rende section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 111 tectonic architecture . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 112 sedimentary units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 112 lønstrup klint formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 112 rubjerg knude formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 112 structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 113 gr01 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 113 gr02 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 114 gr03 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 114 gr04 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 114 gr05 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 115 gr06 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 115 gr07 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 116 gr08 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 116 gr09 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 116 gr10 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 116 gr11 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 117 gr12 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 117 gr13 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 117 interpretation of structural development . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 117 rubjerg knude fyr section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 118 tectonic architecture . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 118 sedimentary units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 121 structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 122 anastomosing thrust-fault brecciation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 122 interpretation of structural development . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 122 stortorn section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 122 tectonic architecture . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 123 st01 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 123 st02 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 123 st03 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 123 st04 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 123 st05 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 124 st06 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 124 st07 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 124 st08 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 124 st09 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 125 st10 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 125 sedimentary units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 126 structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 126 interpretation of structural development . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 126 moserende section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 128 tectonic architecture . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 128 mr01 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 129 mr02 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 130 mr03 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 130 mr04 and mr05 thrust sheets . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 130 mr06–mr08 thrust sheets . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 131 6 mr09 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 132 mr10 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 132 mr11 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 132 mr12 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 132 mr13 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 133 sedimentary units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 134 lønstrup klint formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 134 rubjerg knude formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 134 structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 134 diapir structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 135 thrust faults . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 135 footwall synclines . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 135 interpretation of structural development . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 136 mårup kirke section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 137 tectonic architecture . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 138 mk01 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 139 mk02–mk04 thrust sheets . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 139 mk05–mk07 thrust sheets . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 139 mk08–mk10 thrust sheets . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 140 mk11–mk20 thrust sheets . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 140 sedimentary units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 141 structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 141 interpretation of structural development . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 141 fault-bend-fold model for duplex units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 141 characterisation of thrust duplex mk11–mk20. . . . . . . . . . . . . . . . . . . . . . . . . . . . 143 discussion of structural development . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 143 ribjerg section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 144 ‘store blå’ and ‘lille blå’. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 144 tectonic architecture . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 145 sedimentary units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 146 skærumhede group . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 146 blå-unconformity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 146 ribjerg formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 146 mid danish till formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 147 vendsyssel formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 147 structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 147 interpretation of glacial geology and stratigraphic development . . . . . . . . . . . . . . . . 147 dynamic development of the thin-skinned thrust faulting . . . . . . . . . . . . . . . . . . . . . . 148 moserende section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 148 moserende section: summary data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 149 stortorn section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 149 stortorn section: summary data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 156 rubjerg knude fyr section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 156 rubjerg knude fyr section: summary data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 156 grønne rende section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 156 grønne rende section: summary data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 157 stenstue rende section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 157 stenstue rende section: summary data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 161 sandrende section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 161 sandrende section: summary data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 163 brede rende section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 163 7 brede rende section: summary data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 166 kramrende section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 166 kramrende section: summary data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 168 martørv bakker section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 168 martørv bakker section: summary data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 170 stensnæs section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 170 stensnæs section: summary data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 172 ulstrup section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 173 ulstrup section: summary data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 174 summary of dynamic development . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 174 discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 178 thrust-fault architecture . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 178 balanced cross-section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 179 thrust brecciation and diapirism . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 180 thrust-fault dynamics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 181 syntectonic deposition . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 181 proglacial and subglacial deformation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 182 glacial geological conditions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 183 conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 185 acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 185 references . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 186 appendix 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 190 thrust-fault terminology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 190 appendix 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 192 specification of photogrammetric work . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 192 8 9 abstract pedersen, s.a.s. 2005: structural analysis of the rubjerg knude glaciotectonic complex, vendsyssel, northern denmark. geological survey of denmark and greenland bulletin 8, 192 pp. the rubjerg knude glaciotectonic complex is a thin-skinned thrust-fault complex that was formed during the advance of the scandinavian ice sheet (30 000 – 26 000 b.p.); it is well exposed in a 6 km long coastal profile bordering the north sea in northern denmark. the glaciotectonic thrust-fault deformation revealed by this cliff section has been subjected to detailed structural analysis based on photogrammetric measurement and construction of a balanced cross-section. thirteen sections are differentiated, characterising the distal to proximal structural development of the complex. the deformation affected three stratigraphic units: the middle weichselian arctic marine stortorn formation, the mainly glaciolacustrine lønstrup klint formation and the dominantly fluvial rubjerg knude formation; these three formations are formally defined herein, together with the skærumhede group which includes the stortorn and lønstrup klint formations. the rubjerg knude formation was deposited on a regional unconformity that caps the lønstrup klint formation and separates pre-tectonic deposits below from syntectonic deposits above. in the distal part of the complex, the thrust-fault architecture is characterised by thin flatlying thrust sheets displaced over the footwall flat of the foreland for a distance of more than 500 m. towards the proximal part of the complex, the dip of the thrust faults increases, and over long stretches they are over-steepened to an upright position. the lowest décollement zone is about 40 m below sea level in the proximal part of the system, and shows a systematic step-wise change to higher levels in a distal (southwards) direction. the structural elements are ramps and flats related to hanging-wall and footwall positions. above upper ramp-hinges, hanging-wall anticlines developed; footwall synclines are typically related to growth-fault sedimentation in syntectonic piggyback basins, represented by the rubjerg knude formation. blocks and slump-sheets constituting parts of the lønstrup klint formation were derived from the tips of up-thrusted thrust sheets and slumped into the basins. mud diapirs are a prominent element in the thrust-fault complex, resulting from mud mobilisation mainly at hanging-wall flats and ramps. shortening during thrust-fault deformation has been calculated as 50%. only about 11% of the initial stratigraphic units subjected to thrust faulting has been lost due to erosion. the thrust-fault deformation was caused by gravity spreading of an advancing ice sheet. overpressured mud-fluid played an important role in stress transmission. the average velocity of thrust-fault displacement is estimated at 2 m per year, which led to compression of a 12 km stretch of flat-lying sediments, c. 40 m in thickness, into a thrust-fault complex 6 km in length. the thrust-fault complex is truncated by a glaciotectonic unconformity, formed when the advancing ice sheet finally overrode the complex. when this ice sheet melted away, a hilland-hole pair was formed, and meltwater deposits derived from a new ice-advance (ne-ice) filled the depression. the ne-ice overran the complex during its advance to the main stationary line situated in the north sea. when this ice in turn melted away (c. 19 000 – 15 000 b.p.), the glacial landscape was draped by arctic marine deposits of the vendsyssel formation (new formation defined herein). _________________________________________________________________________________________ author’s address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: sasp@geus.dk 10 55n 5e 10e 15n 60n scandinavian ice sheet norway denmark copenhagen göteborg møns klint bovbjerg fur knudeklint hanklint mols hoved rügen ristinge klint lønstrup klint sweden germany baltic ice advance swedish ice advance norwegian ice advance 28 ka bp 30 ka bp 20 ka bp 17 ka bp 17 ka bp 17 ka bp 28 ka bp fig. 1. map of the danish basin indicating the distribution of the scandinavian ice sheet during the three main ice advance events, with source areas in southern norway, central sweden and the baltic, in the middle–late weichselian. the approximate timing of the stationary lines are given; the early progressive ice advance is indicated in black, the subsequent late ice border lines in red. the locations of major glaciotectonic complexes formed during the ice advances are indicated by asterisks. 11 introduction glaciotectonic studies in denmark have a long tradition, and an important part of structural geology studies in denmark concern glacial tectonic deformation resulting from the southward advance of the scandinavian ice sheet in the pleistocene (fig. 1). the description of the geological structures dates back to puggaard (1851), who made one of the first extensive and detailed danish structural analyses of a tectonic complex and provided a classic cross-section of møns klint. johnstrup (1874) established the concept of glacial deformation. the next milestone in glacial tectonic studies in denmark was by jessen (1918, 1931), whose detailed survey of lønstrup klint (fig. 1) included a structural analysis and an attempt at a glaciodynamic interpretation of the deformation structures observed. the lønstrup klint coastal section includes the rubjerg knude glaciotectonic complex, which is the subject of this study (fig. 2). a danish school of glaciotectonic studies subsequently developed (madsen 1916; jessen 1931; gry 1940, 1941; rosenkrantz 1944; berthelsen 1973, 1975, 1978, 1979; sjørring 1974, 1977, 1981, 1983; rasmussen 1975; petersen 1978; houmark-nielsen 1987, 1988; pedersen 1987, 1993, 1996, 2000; pedersen & petersen 1988, 1995, 1997; pedersen et al.1988; klint & pedersen 1995; jakobsen 1996), which has naturally been stimulated by geologists working with glaciotectonic structural geology internationally (banham 1977, 1988; stephan 1980; aber 1982, 1993; ehlers 1983; van der wateren 1985, 1992; boulton 1986; boulton & hindmarsh 1987; croot 1987, 1988; meer 1987; goldthwait & matsch 1988; aber et al. 1989; hart 1990; hart & watts 1997; bennett 2001). the similarity in structural geometry between glaciotectonic terrains and orogenic belts has led to prolonged debate. are glaciotectonic terrains scale models for orogenic deformation? or does the soft and synsedimentary nature of glaciotectonics differ in principle from that of fold belt deformation? arguments for deformational similarity have been put forward by berthelsen (1978, 1979), banham (1988), aber et al. (1989), van der wateren (1992) and pedersen (1987, 2000). these structural geologists share the opinion that the terminology of structural geology related to orogenic belts is applicable in the description and discussion of glaciotectonic complexes. the main differences between deformation in metamorphically altered rocks and glaciotectonic deformation of soft sediments are: (1) the presence of ‘free’ water, which enables liquefaction and fluidisation, (2) the velocity of the deformation, and (3) the shallowness of penetrative deformation. in contrast, deformation of metamorphic rocks commonly involves alteration and recrystallisation of minerals, processes that never apply to glaciotectonics. the advantage of a study of glaciotectonic complexes is that the structures are at a scale that allows them to be studied in a single exposure, in contrast to fold belts where extensive field mapping and expensive geophysical investigations are typically required for adequate documentation of the structures. furthermore, many glaciotectonic complexes are geologically young, which means that the upper structural levels are still preserved and interpretation of the full dynamic development of structural complexes is possible. the structural architecture of glaciotectonic complexes may therefore serve as inspiration for the interpretation of thin-skinned structural relationships in fold belts and thrust-fault deformation terrains. the structural analysis of the rubjerg knude glaciotectonic complex is presented as a mesoscopic model of a thinskinned thrust-fault complex (plates 1, 2). history of the present investigation this study focuses on the structural framework and dynamic development of the glacial tectonic thrustfault complex at rubjerg knude, lønstrup klint. it is based on twenty years of investigations of the lønstrup klint cliff section. the author took up the study of glacial tectonic thrust-fault structures after having concluded a ph.d. thesis on thin-skinned thrust faulting in the north greenland fold belt (pedersen 1979, 1981, 1982, 1986a, 1987). a large part of the study of the fold belt structures in peary land, north greenland, was photogrammetric mapping (pedersen 1979, 1981), undertaken at a time when geological mapping by computer-assisted photogrammetry was under development in copenhagen. this project was an integrated collaboration between the geological survey of greenland, the institute of surveying and photogrammetry of the technical university of denmark (dtu), the geological museum (gm) and the geological institute (gi) of the university of copenhagen. in the years 12 l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l diamictite glaciolacustrine and glaciofluvial sand and gravel mobilised mud non-marine sand non-marine clay and silt marine clay rubjerg knude formation lønstrup klint formation stortorn formation 0 500 1000 1500 2000 2500 3000 0 100 200 300 400 5 150014001300120011001000 2000 2100 2200 2300 2400 2500 350034003300320031003000 4100 4200 4300 4400 4500 560055005400530052005100 steps ribjerg moserende kramrende stenstue rende stortorn grønne rende brede rende stensnæs fig. 2. geological cross-section of the rubjerg knude glaciotectonic complex. for details and legend, see plate 1. 13 l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l aeolian dunes holocene peat marine clay and sand sandy till glaciofluvial sand { mid danish till formation & kattegat till formation vendsyssel formation ribjerg formation top of dunes clifftop – glacial abrasion surface thrust, fault unconformity intraformational bedding 3500 4000 4500 5000 5500 6000 m 500 600 700 800 900 1000 m 2000 m1900180017001600 2600 2700 2800 2900 3000 m 4100 m40003900380037003600 4600 4700 4800 4900 5000 5100 m 6000 m590058005700 steps rubjerg knude fyr sandrende martørv bakkeroddervej ulstrup ulstrup rende tvonnet rende mårup kirke 14 up to 1990, techniques of geological mapping and construction of geological cross-sections based on multimodel photogrammetric analysis were developed and made available at dtu (dueholm 1992). initial investigations in co-operation with k. dueholm (dtu) and a.k. pedersen (gm) proved the applicability of multimodel photogrammetry in the study of glaciotectonic cross-sections in denmark by an examination of the møns klint clif f section (pedersen 2000). subsequently, the photogrammetric investigation of the rubjerg knude clif f section was initiated, and forms the basis of the present work. objectives the objectives of the study of the rubjerg knude glaciotectonic complex can be summarised as follows. 1. a description of an exceptionally well-exposed glaciotectonic complex, which can be taken as an example of a very low friction thrust-fault wedge, presented as a detailed cross-section based on multimodel photogrammetric measurements of the rubjerg knude clif f section. 2. a demonstration of the techniques of balanced cross-section construction that permit interpretation of the unexposed parts of the thrust-fault complex. 3. the construction of a model for the dynamic development of the proglacial thrust system that demonstrates the sequential evolution of increasing deformation intensity and the interplay with syntectonic depositional processes. 4. an interpretation of deformation processes within the framework of danish glacial stratigraphy in the late pleistocene (late middle to late weichselian c. 30 000 – 20 000 years b.p.). previous conceptual models the basic concept of glacial processes acting as the deformation agent was formulated by johnstrup (1874). his concept was primarily focused on the formation of the spectacular cliffs at møns klint in south-eastern denmark and on rügen in north-eastern germany. however, subsequently johnstrup (1882) also included the formation of the steeply inclined floes exposed in the lønstrup klint clif f section in the classic examples of glacial deformation in denmark. (the term floes is frequently used in the old glacial geology literature inspired by the idea that the dislocated sheets were groundor permafrozen; in a structural geological context, floes are identical to thrust sheets or thrustsheet segments.) johnstrup’s main conclusions concerning the glaciotectonic origin of the deformation at lønstrup klint were: (1) the dislocations are superficial without extending down to a deep root zone, and are restricted to surface phenomena, (2) the direction of movement indicated from the dip of the dislocated floes corresponds to a uniform direction of ice advance, and (3) the dislocated floes formerly constituted one undisturbed area. the detailed mapping and construction of the cross-section was presented by jessen (1918) in his geological description of the vendsyssel map sheet. however, the final detailed description of the dislocations at lønstrup klint was published later (jessen 1931). in 1927, george slater included a study of the lønstrup klint section as part of his thesis for a d.sc. degree at the university of london, which also included a study of glacial deformation at møns klint. the most striking conclusion was that the glacial deformation at lønstrup klint was caused by englacial deformation. slater (1927, p. 312) summarised thus: “… 2. the deposits represent the final positions of englacial material after the melting of the interstitial ice. 3. the type of structure is analogous to that seen in decaying arctic glaciers, and is due to the arresting of movement of the frontal part of an overloaded ice-sheet. 4. the structure has been built up in the reverse direction to the line of movement.” slater (1927) interpreted the lønstrup klint section as a variety of glacial tectonics he termed ‘the stagnant-glacier type’. subsequently, axel jessen and karl gripp exchanged ideas about proglacially formed glaciotectonic structures, and concluded that the structures jessen had observed at lønstrup klint were similar to those that gripp (1929) described from the foreland of the adglacial tectonics – concepts and models 15 vancing holmströms gletscher on spitsbergen. in his detailed and comprehensive description of his investigations, jessen (1931) concluded that the disloctions cannot have formed englacially, but must be the result of pressure building up due to loading at the margin of the advancing ice. this pressure spreads out laterally into the clayey units, which in the foreland react by splitting up into fractured dislocation sheets compressed in front of the advancing ice masses. jessen (1931) also discussed the dif ficulty related to the displacement of the sheets without fracturing of the lithological units resulting in a complete collapse during deformation, and he pointed out that johnstrup (1882) had suggested that the deformed layers could have been ground-frozen. jessen’s (1931) more subjective arguments against slater’s work concern the fact that slater (1927) did not refer to jessen’s (1918) substantial work on vendsyssel and in particular his published cross-section of lønstrup klint. jessen pointed out that major anticlines in slater’s crosssection between mårup kirke and rubjerg knude fyr do not exist, and that slater’s (1927) misinterpretation must be ascribed to his superficial investigations which did not allow him to check the way-up relationship of each limb in the fold structure (jessen 1931). in his work on the glaciotectonic deformation of palaeogene diatomites with ash layers in the limfjorden region, gry (1940) compared these with the deformation at lønstrup klint and supported the proglacial deformation concept of gripp (1929) and jessen (1931). furthermore, gry proposed a gravity-spreading model for the deformation and attempted a very early balanced cross-section in the consideration of restoration of the dislocated thrust sheets (fig. 3). however, gry (1940) proposed a cylindrical model for the thrust surfaces, and in his ‘back-stripping’ cross-section the floes were displaced along circular fault lines. thus, in his dynamic consideration the floes were assigned a standing position with their frontal parts ‘up in the air’ (fig. 3), and he consequently concluded that more than 80% of the upper sand-series at lønstrup had been eroded away by the advancing ice. in contrast to this point of view, pedersen (1987) suggested that a large proportion of the upper sandseries was deposited syntectonically; this removed the requirement that a large part of the floes or thrust sheets had been eroded away. pedersen (1987) interpreted the glaciotectonic thrust-fault complex as an example of gravity-spreading deformation, viewed in the light of the gravity-spreading experimental model presented by bucher (1956) and with reference to comparable gravity-spreading deformation in soft sedimentary rocks exemplified by the mudlumps in the mississippi delta (morgan et al. 1968). furthermore, the mudlumps or mud diapirs in the lønstrup klint imbricate fan were described, and interpreted as an integral part of a conceptual dynamic model for thrust-fault related mud diapirism and syntectonic sedimentation (fig. 4). sadolin et al. (1997) elaborated on the model of syntectonic sedimentation in the lønstrup klint section. based on detailed sedimentological studies, they pointed out the importance of the unconformity that separates the lower muddy units (their unit a), from diluvial sand yoldia clay fig. 3. a model for structural balancing of the dislocated floes in the lønstrup klint section suggested by gry (1941). in his model, the displacement surfaces were regarded as cylindrical sections and due to the suggested amount of displacement about 80% of the dislocated floes was subsequently eroded away. 16 the upper sandy units (their units b–d). the lower unit a was interpreted to have been deposited in a lake isolated from the former marine kattegat–skagerrak basin by either a damming of the advancing ice, in accordance with ideas also presented by jessen (1918, 1931), or simply by isolation of the lake basin due to lowering of sea level in the late pleistocene (sadolin et al. 1997). the unconformity was interpreted to reflect a major drainage event of the lake basin before a shallow lacustrine basin was established, characterised by incursions of glaciofluvial deposition (units b–d of sadolin et al.1997). during the deposition of units c and d, glaciotectonic thrusting commenced contemporaneously with the rise of mud diapirs and the formation of normal faults due to mass adjustments in the mobilised mud in the subsurface (sadolin et al. 1997; fig. 5). the conceptual model presented here aims at an interpretation based on the concepts of thin-skinned thrust-fault tectonics. although the scale is an order of magnitude smaller than in typical orogenic belts, it has not been found appropriate to introduce special terminology for the deformation structures in the rubjerg knude glaciotectonic complex. the concept of thrust-fault deformation and related structures is summarised in the following chapter. thin-skinned thrust faulting: the concept it is difficult to judge exactly when the concept of thin-skinned thrust faulting nucleated, as it represents a gradual evolution of ideas over the last 25 years or more. however, boyer & elliot (1982) appear to have been the first to give a conceptual introduction to the basic principle of thin-skinned thrust faulting. suppe (1983, 1985) improved the concept by defining and describing the geometry and kinematics of fault-bend folding. jamison (1987) and schirmer (1988) contributed with further improvements of geometric analysis of fold development in overthrust terranes and thrust-fault hanging-wall successions. mcclay (1992) presented a glossary of thrust tectonic terms, and erickson & jamison (1995) demonstrated viscous-plastic finite-element models of fault-bend folds. in 1997, an entire volume of the journal of structural geology was devoted to thrust-fault tectonics. among the papers that particularly inspired and supported this study of glaciotectonic thrust faulting were those of contreras & sutter (1997), medwedeff & suppe (1997) and mitra & sussman (1997). thrust-fault modelling to better understand the range of possible configurations of different structural frameworks of thrust-fault complexes, a series of computer models were tested with the aid of the program autofault, a ‘balanced cross section program’ within the autocad system frame (ozkaya 1994). four of these test models are demonstrated here to illustrate the thin-skinned thrustfault concept (figs 6–9). the basic function of the model is to define and construct a layer package onto which a thrust fault is added and given a certain displacement. the program then calculates the configuration of the thrust sheet fig. 4. a four-stage model for the development of mud diapirs related to thrust faulting in lønstrup klint suggested by pedersen (1987). note that in the model the thrust zone of the hanging-wall ramp constitutes mobilised mud and that syntectonic deposits accumulate ‘piggyback’ between the thrust sheets. 17 100 m 100 m 100 m 100 m sandrende thrust fault c c c b b b b a a a a a b c a mud diapir x1 a x2 a c c mobilized mud y1 y2 extensional faults mobilized mud x1 x2 y1 y2 x1 x2 y1 y2 position of thrust-fault fractures unconformityerosionally removed 2 1 d d m. s. l. older yoldia clay 4 3 fig. 5. the structural and depositional development of the sandrende section suggested by sadolin et al. (1997). the model summarises four stages of development initiating with the formation of the regional erosional unconformity (1). unit b was deposited in topographic lows above the unconformity, and thrust faulting initiated contemporaneously with the deposition of unit c (x 1 –y 1 and x 2 –y 2 denote same reference points separated by the thrusts, where x = footwall syncline and y = hanging-wall anticline) (2). propagation along the thrust faults continued and unit c was deposited during increasing tilting of the thrust sheet. normal-fault fractures formed in connection with the incipient diapirism (3). the sandrende diapir rose during deposition of unit d and normal faulting propagated. in the proximal part of the thrust sheet, a network of conjugate extensional faults developed and interference between a new-formed satellite thrust and the normal faults af fected the complex. the tip of the thrust sheet was bent due to drag along the side of the rising diapir (4). star symbol provides a reference point through the development stages. 18 step 1 step 2 step 3 step 4 step 5 step 6 hanging-wall block ramp footwall block 50 m displacement 100 m displacement 150 m displacement 200 m displacement 300 m displacement 400 m displacement lower flat axial surface upper flat = top surface lower ramp hinge upper ramp hinge hanging-wall anticline hinterland-dipping limb foreland-dipping limb hanging-wall flat hanging-wall flat hanging-wall ramp upper footwall flat footwall ra mp décollement or lower footwall flat 0 100 200 300 400 m n s fig. 6. test model 1 of thrust-fault deformation constructed with the computer program autofault (ozkaya 1994). the model demonstrates the development of simple ramp propagation given increasing displacements. in the first four steps, the displacement is sequentially increased by 50 m, whereas a displacement of 100 m is added to steps 5 and 6. note that a ‘typical upright anticline’ develops when the displacement is about twice the thickness of the layer package displaced. moreover, the model illustrates the terminology applied in the text and defined in appendix 1. 19 for the specific model constructed. thus the program gives the ‘differential’ calculation model to an induced ‘integration’ solution configuration. further thrust faults can be added, and be given new displacements, such that rather complex models can be constructed. however, a few limitations of the program hamper realistic comparisons with nature. thus the program cannot handle inclinations exceeding 60°. in general this is not a problem as ramp angles typically range between 10° and 35° and for rock mechanical reasons never exceed 45° (ozkaya 1994). however, the problem of steep inclinations becomes important in complexes including superimposed deformation. a second limitation is that testing with superimposed displacements requires a construction with an upper flat located within the model. this results in an unrealistically high number of shallow upper flats in the models, as illustrated below in test model 4 (see fig. 9). thirdly, the program cannot accommodate cross-cutting thrust-fault relationships, which limits the spacing and dip of ramps. nevertheless, the test models give a good introduction to the thrust-fault concept, and demonstration of models with basic layer package dimensions approaching the scale of thrust sheets involved in the rubjerg knude glaciotectonic complex can be achieved. a glossary of the thrust-fault terms used here is given in appendix 1; note that only contractional thrust-fault structures are considered. test model 1 the first autofault model displays a simple thrust fault with one ramp connecting a lower and an upper flat (fig. 6). the development of thrust-fault structures, in particular the fault-bend folding of the hangingwall anticline, is demonstrated in six steps with increasing displacement. the ramp angle is 25°, and the layer package constitutes a lower unit 25 m thick where the lower flat (or the décollement zone) is located. above this, one 25 m and two 20 m thick layers have been constructed, with a 30 m thick uppermost layer (fig. 6). the model approaches the assumptions of parallel behaviour with preservation of layer thickness, no net distortion where layers are horizontal, and conservation of bed length (suppe 1983). step 1 shows the gentle hanging-wall anticlinal folding after 50 m displacement. note the flat-topped nature of the hanging-wall anticline, which makes it almost insignificant. the backlimb of the anticline dips toward the left, parallel to the ramp, and the axial surfaces defined by the bend above the lower ramp hinge and the bend of the hanging-wall anticline define two kink bands dipping steeply to the right. by comparing steps 1 and 2 it can be seen that the spacing between the kink bands increases with increasing displacement. step 2 gives the configuration after 100 m displacement. here the forelimb dipping towards the foreland to the right starts to be a significant part of the structure. note the increase in spacing between the kink bands in the backlimb structure. the kink bands define minor zones of weakness, which could develop into small reverse faults as in the thrust model demonstrated by wiltschko (1979). these are referred to as back thrusts. step 3 shows the structural development after 150 m displacement. note that the flat-topped hangingwall anticline now has a more angular upright form, where the kink bands fanning up from the positions near the upper ramp hinge approach each other. however, in the model the anticline maintains its flat-topped structure and retains two axial surfaces (kink bands). step 4 demonstrates the formation of the upright, angular hanging-wall anticline, where the amount of displacement is close to the length of the thrust-fault ramp. due to the geometric adjustments the hangingwall ramp is shorter than the footwall ramp. the displacement is 200 m corresponding to about two times the thickness of the thrust sheet. step 5 shows the structural development after 300 m displacement. the hanging-wall anticline becomes even more flat-topped and the space between its axial surface kink bands increases. note that the forelanddipping forelimb is linked to the hanging-wall ramp displaced along the footwall flat, and the hinterlanddipping backlimb corresponds to the hanging-wall flat bent up along the footwall ramp. step 6, with a displacement of 400 m demonstrates that the main structural configuration is maintained, except for the increase in spacing between the backlimb and the forelimb. test model 2 the second autofault model demonstrates the propagation along a thrust fault dif ferentiated into a décollement zone, a lower ramp, an intermediate flat, an upper ramp and an upper flat bringing the thrust fault up to the top surface (fig. 7). the model is con20 structed with two lower units, 40 m in thickness; the décollement zone is located in the second layer. the lower layers mimic the lower clay units of the lønstrup klint stratigraphy, and two c. 25 m thick layers overlie them. the top layer is 50 m thick, but while not comparable to any part of the stratigraphy in the lønstrup klint section, its construction yields a better demonstration of the development envisaged. the lower ramp is given a dip of 25° and the upper ramp a dip of only 15° to reflect the principle of increasing angle of fracturing with increasing depth (hobbs et al. 1976; pedersen 1996). the distance between lower and upper ramps along the intermediate flat is c. 250 m, and three steps are presented in fig. 7. step 1 is given 50 m displacement and two hanging-wall anticlines immediately appear. the steep ramp clearly initiates the formation of an upright anticline with steeply dipping limbs. between the two hanging-wall anticlines, an intervening syncline forms above the intermediate flat. the involute surface of the syncline provides the location for a broad, shallow basin. step 2 shows the structural development after 100 m displacement. this demonstrates clearly that the intervening syncline is an obvious site for a piggyback basin to develop. note that the steeply dipping forelimb of the hanging-wall anticline above the lower ramp would be the obvious site for erosion and the source of material feeding into the piggyback basin. step 3 demonstrates that with a displacement of 200 m, the piggyback basin becomes narrow and is elevated to a higher position as a consequence of the displacement up along the upper ramp; it is eventually lifted out of the position for being a centre of deposition. with increasing displacement, the frontal part of the thrust sheet develops into a wedge-shape structure. hanging-wall block lower hanging-wall ramp lower ramp upper ramp upper flat upper hanging-wall ramp upper hanging-wall ramp footwall block 50 m displacement 100 m displacement 200 m displacement piggyback basin piggyback basin intermediate flat lower flat fault-bend folding step 1 step 2 step 3 0 100 200 m n s fig. 7. test model 2 of thrust-fault deformation constructed with the computer program autofault. the model demonstrates the development of thrust-fault propagation along a lower and an upper ramp and the connecting flats. note in this model the formation of two anticlines divided by a syncline, the depression of which is the obvious location of a piggyback basin. 21 test model 3 the third autofault model aims at constructing an imbricate complex by branching faults fanning up from the same décollement level (fig. 8). the model is constructed with a lower 20 m thick unit in the top of which the décollement zone is located. above the décollement zone, three units with a combined thickness of 50 m form the lower part of the thrust sheets, and the succession is capped by an upper 20 m thick unit. in three sequential steps, the principle of piggyback thrusting is demonstrated (fig. 8). step 1 shows 100 m displacement along a deeprooted ramp dipping 30°. note the normal architecture of the hanging-wall anticline results from the ramping (compare with fig. 6, step 3). step 2 demonstrates the re-orientation of the piggyback thrust sheet by the introduction of 100 m displacement along a 18° dipping ramp in front of and below the first thrust fault. note that the accumulated displacement of the first thrust sheet amounts to c. 200 m. step 3 shows an additional 100 m displacement along a low-angle 12° dipping ramp. although the model demonstrates the main architecture of the imbricate fan illustrated by pedersen (1987), it is a fairly simple model which may have only little relevance to natural conditions. test model 4 the final autofault model demonstrates the more likely formation of a steeply dipping imbricate fan or duplex (fig. 9). the model is given the same stratigraphic units as in test model 3 (fig. 8). a longer décollement zone is located in the middle of the lowermost unit, in addition to an intermediate flat in the third layer, while the upper flats are located within the uppermost unit. the initial steps in the construc1 1 2 3 1 2 100 m displacement 100 + 100 m displacement 300 m accumulated displacement simple ramp piggyback thrust sheet branching thrust fault branching thrust-fault imbricate fan step 1 step 2 step 3 0 100 200 m n s fig. 8. test model 3 of thrust-fault deformation constructed with the computer program autofault. the model demonstrates the formation of an imbricate fan by successive thrust-fault splays branching up from the main décollement zone. the encircled numbers refer to the sequential phase of thrust imbrication. the model is probably not comparable to structures formed in nature, but can be regarded as an introduction to test model 4 (fig. 9). 22 tion of this model are similar to the examples demonstrated above, and hence only the final two steps are illustrated (fig. 9). however, these give a convincing illustration of the increase of dips in an imbricate thrustfault complex. step 1 illustrates the final structural architecture after 140 m displacement of thrust sheet 1 along the décollement zone, the lower ramp, the intermediate flat, an upper ramp and onto the upper flat (dips of ramps c. 25°). thrust sheets 2–5 were formed by branching ramps (dip of ramps c. 15°) with a displacement of c. 80 m added to each thrust fault. finally, the leading thrust sheet (6) is displaced 90 m along the lower décollement zone and a deep-rooted 30° dipping ramp. note that the branching ramp imbricates are carried piggyback on thrust sheet 6. furthermore, it should be noted that a long trailing segment of thrust sheet 6 occurs between the décollement zone and the intermediate flat. if this trailing segment becomes chopped up into duplexes between the two deep-rooted ramps, it will affect the overlying imbricates by vertical elevation and the formation of antiformal stacks. step 2 illustrates the over-steepening of the imbricates stacked onto the backlimb of the hanging-wall anticline of thrust sheet 6 arising from the addition of 100 m displacement to step 1 along the leading thrust rooting down to the lower décollement zone. test models: concluding remarks a set of principles may be derived from the test models. 1. the level of elevation of the reference surface is directly related to the number and sizes of ramps the thrust sheet has passed. a ramp rooting down to a deep flat level corresponds to a high elevation of the topmost reference surface. in contrast, if a top reference surface is positioned at the same level as in the foreland, the thrusting corresponds to a translation along a flat. 2. the steeper the ramp, the earlier its time of formation. gently dipping ramps are initiated at a late stage of deformation in areas proximal to the foreland. 3. the thickness of a piggyback basin reflects its duration as depocentre. thus a small thickness of piggyback basin fill indicates an early trapping of the basin by overthrusting of a hanging-wall block. 4. a thick succession in the piggyback basin reflects a long period of translation of the thrust sheet along a long flat. step 1 step 2 fig. 9. test model 4 of thrustfault deformation constructed with the computer program autofault. the model demonstrates an imbricate fan (see fig. 8) subjected to faultbend folding during piggyback translation of an underlying hanging-wall flat propagation along a footwall ramp. the footwall ramp propagation will consequently result in increasing dips of the thrust sheets in the imbricate fan. encircled numbers indicate successive thrust sheets. concept of balanced cross-section the principle of the balanced cross-section in structural analysis of thrust-fault systems was elegantly outlined by dahlström (1969) and further improved by suppe (1985). the application of balanced crosssections in glaciotectonics has been demonstrated by croot (1987), klint & pedersen (1995) and pedersen (1996). in the construction of the balanced section, two different functions are applied: (1) the line balance, and (2) the volume balance, which in a 2-d crosssection corresponds to area balance. the first function concerns the length of displacement, whereas the second function concerns the preservation of volume in the deformed cross-section compared with the restored undeformed cross-section (for demonstration see plate 2). the basic method of balancing a crosssection (dahlström 1969) is restoration by defining a pinpoint to be fixed to the foreland and then restoring the thrust sheets back to their initial pre-deformational position. thus one begins at the foreland and then by line balancing the thrust sheets are pulled back sequentially to their position prior to displacement. this requires a measure of displacement, which is the essential, but often difficult figure to achieve without some range of uncertainty. details concerning the construction of the balanced cross-section of the rubjerg knude glaciotectonic complex (plate 2) are given below. 23 location the rubjerg knude cross-section is 6124 m long and extends from the coastal cliff immediately south of lønstrup, ribjerg, to about 300 m north of the ramp leading down to the beach at nørre lyngby (fig. 2, plate 1). the strike of the section is 17°, which is nearly parallel to the direction of the coastline. this is also approximately perpendicular to the main concentration of structural strikes (bedding, thrust faults and fold axes; fig. 10). the cross-section was consequently constructed to fit a general plane of orthographic projection with a projection axis striking 107°. the rubjerg knude cross-section covers only the rubjerg knude glaciotectonic complex. thus it is not as extensive as the cross-section of lønstrup klint constructed by jessen (1918, 1931), which extends from the cliff at the northern fringe of lønstrup to the northern part of the beach at løkken (see fig. 12). the utm co-ordinates (zone 32, ed50) of the end points of the rubjerg knude cross-section are 547512, 6370243 (n-end point) and 545251, 6364783 (s-end point). photogrammetric work the cross-section of rubjerg knude glaciotectonic complex (plate 1) is based on a multi-model photogrammetric investigation of the cliff section using the method described by dueholm (1992). oblique photographs were taken from a cessna fixed-wing aircraft with a minolta xg2 camera with known optical specifications, calibrated at the laboratory of photogrammetry at the technical university of denmark. standard 24 × 36 mm diapositive colour film was used, and the photographs were taken with 66% overlap from a distance of 200–300 m with an inclination angle of c. 35°, which provided the basis for setting up 67 stereoscopic models. in the laboratory, the orientation of the stereo-models was carried out based on ground control points adapted from two sets of vertical aerial photographs at a scale of 1:25 000, namely d9202 g 1365–66 and kms 9203 a509–10 taken in may 1992. the stereoscopic instrument used was a kern dsr 15 analytic plotter with a dec vms operating system and the special attached geoprogram developed n s 0 100 200 m 100 m displacement on youngest thrust fault 200 m displacement on youngest thrust fault 1 1 2 3 5 2 3 4 5 4 6 6 location and construction of the rubjerg knude cross-section 24 1 2 3 n n = 52 1 2 3 n = 60 n 1 2 3 n = 83 n 1 2 3 n = 83 n a b c d fig. 10. stereographic projection diagrams of the orientation of structural elements in the rubjerg knude cross-section. the stereograms, lower hemisphere, equal area (schmidt) net, display the concentration of the poles to bedding planes (black dots) or thrust planes (black triangles). a and b are measurements taken from jessen (1931), and c and d are data produced in this study. contour intervals are 1, 2.5, 5, 7.5, 10, 12.5, and 15%. the density point in all four diagrams is close to 197°/35°. comparing the two sets of diagrams demonstrates that the structural orientation has been maintained despite c. 100 years erosion corresponding to c. 125 m retreat of the coastal cliff section. black squares (d) indicate normal fault planes. blue lines/numbers indicate principal compression axes. 25 by dueholm (1992). in the stereoscopic models, the geological features were outlined by the floating mark and digitised by the attached computer. the digitised data were stored for the later construction of the crosssection and the transformation for other programs applied for the management of the cross-section display. the scale of the rubjerg knude cross-section in the analytic plotter version is 1:500, and the accuracy of the plotted data is about 25 cm (for further details, see appendix 2). digital editing in order to represent the cross-section in a publishable display, the digitised data were transferred to arcinfo at the gis-laboratory at the geological survey. here it was transformed into an arc-view project, which served as the computer tool for editing the crosssection. thus all areas were converted to closed polygons, which were annotated to fit the legend of lithologies. during this editing, interpretations were made to finish the display of the cross-section, in particular interpretations of the scree-covered parts of the section. this was carried out contemporaneously with the construction of the balanced cross-section (see below), and a few additional corrections were added to the rubjerg knude cross-section. some new exposures along the cliff section appeared in 1997–1999, which added to a better understanding of the structures in the transition from the frontal part of the glaciotectonic complex to its foreland. these have been incorporated into the arc-view project. the final editing of the cross-section concerned the balanced cross-section. the construction of the balanced section was digitised and transformed into an arc-view project, and the subsequent interpretation of the extension of the thrust-fault ramps below sea level was added. thus the rubjerg knude cross-section comprises a display of the exposed part of the cliff section with lithological and structural identity added as themes. furthermore, the cross-section includes an interpretation of the thrust-fault structures in the subsurface. finally, a balanced construction was added l l 2900 3000 m dc dm ds α α l/r-u l/r-u t t t t l/r-u l/r-u fig. 11. illustration of the method used for estimation of the displacement for the balanced cross-section. above the main erosional unconformity at the top of the cliff, the extension of the thrust sheet tip is constructed by the intersection between the thrust fault (t) and the l/r-unconformity (l/r-u) based on the angle (±) between the bedding of the thrust sheet and the hanging-wall ramp. dm, displacement measured; dc, displacement constructed from tip-extension; ds , displacement estimated from the interpretation of thrust-fault trace under the scree cover. the section illustrated is part of the rubjerg knude fyr section (plate 1). 26 table 1. the distribution of areas in the balanced cross-section (plate 2) balance (plate 2a) ramps (plate 2b) section* number of areas area (m2) section* number of areas area (m2) 01ul 5 23 048 01ul 13 24 302 02sn 13 8965 02sn 18 8536 03mb 15 28 443 03mb 21 30 944 04kr 10 24 158 04kr 22 18 390 05br 28 34 143 05br 40 33 548 06sr 28 33 218 06sr 49 31 588 07ss 32 26 421 07ss 31 23 973 08gr 55 49 842 08gr 47 45 118 09rf 30 22 827 09rf 26 21 458 10st 54 43 674 10st 41 36 656 11mr 69 51 902 11mr 55 45 342 12mk 95 82 226 12mk 87 62 763 13bl 8 17 922 13bl 2 14 313 nrly 2 5437 13ri 1 4405 ptr 3 2538 md 1 472 ve 4 9818 * the annotated numbers of sections (05br) correspond to the sequential location of each section in a distal–proximal order, and the capitalised letters refer to the general abbreviation of the section names (see plate 2). to the cross-section project, such that each thrust sheet is annotated in a balanced restored cross-section as well as in the structural cross-section displaying the geometry of the ramps and flats (plate 2). construction of the balanced cross-section the construction of the balanced cross-section for the rubjerg knude glaciotectonic complex was based on the geological cross-section, which displays the geometry of the thrust sheets in sufficient detail to allow calculations of their displacements and cross-sectional areas (plates 1, 2). the method of balancing necessitates that the thrust sheet closest to the foreland is the first to be restored to its pre-deformational position. therefore, the balancing works backwards from the distal to the proximal deformation area, and consequently the annotation of the thrust sheets begins with the first thrust sheet restored. in the balanced cross-section of the rubjerg knude glaciotectonic complex, the thrust sheets are additionally annotated according to that part of the cliff in which they occur: two capital letters refer to the name of the section and a number refers to its position from leading edge to trailing end of the section. thus, kr01 is the thrust sheet nearest to the foreland in the kramrende section. a thrust fault is referred to according to the thrust sheet it displaces. however, the trailing footwall ramp is referred to the annotation of the footwall block, which underlies the hanging-wall ramp/flat of the thrust sheet displaced over it. thus the kr02 hanging-wall ramp is displaced up along the kr01 footwall ramp. although one of the basic conditions in constructing balanced sections is the preservation of volumes, which in the areas strongly affected by mud remobili27 sation and diapirism is difficult to maintain, the exercise has been carried out to match a balanced section to the mapped and interpreted thrust-fault framework. so despite the uncertainties and the demand for interpretation of the geometry and magnitude of eroded thrust sheet tapers, the construction of the balanced section added significantly to the understanding of the duplex framework (plate 2b). in the rubjerg knude cross-section (plate 1), the displacement is measured and estimated mainly from the distance between the intersection of the l/r-unconformity (the unconformity between the lønstrup klint and rubjerg knude formations, see below) and the footwall ramp, and the intersection of the l/runconformity and the hanging-wall ramp (fig. 11). however, the tips of the thrust sheets are generally eroded away, so the first approximation is from the l/r-unconformity footwall point to the point where the hanging-wall ramp is truncated by the glaciotectonic unconformity at the top of the cliff. the second approximation is the addition of the distance estimated from the size of the tip eroded away. this estimate is based on a simple geometric construction of the tip-triangle from the dips of the hanging-wall ramp and the l/runconformity, respectively (fig. 11). this line balance is subsequently controlled by the width of the piggyback basin more or less corresponding to the upper footwall flat. all the measured displacements are strictly restricted to the minimum distance to avoid unrealistic exaggerations. therefore the actual displacements might be slightly greater. the area balance is based on a calculation of all the areas annotated in plate 2. the computer-supported calculation was carried out with the arc-info program, and the calculations of the areas in the balanced cross-section and the ramp cross-section deviate by less than 10% (plate 2a, b). this is regarded as a reasonable correspondence considering the various sources of error (table 1). in general, the sections have a smaller area in the ramp cross-section (plate 2b) due to the erosion of areas above the main headof-cliff unconformity, and in most sections the number of areas is higher due to the increased complexity of the geometry in the reconstructed structural cross-section (plate 2b). geological setting the rubjerg knude glaciotectonic complex incorporates deformed sedimentary deposits that belong to the upper part of the mainly marine succession known previously as the skærumhede series (jessen et al. 1910). this succession was deposited in the northern part of the danish basin in the late pleistocene, after the late saalian terrestrial glaciation retreated from denmark (houmark-nielsen 1987, 1999; knudsen 1994). the major source area for deposits in this part of the danish basin is the scandinavian basement in southern norway and central sweden, that comprises precambrian fennoscandian granites and gneisses overlain by palaeozoic metasediments, including permian volcanics and their related intrusive magmatic rocks of the oslo province (oftedahl 1981). the extrabasinal indicator boulders reflect these source areas, which were situated between the centres of ice-cap nucleation and the depositional basin (milthers 1909; smed 1995). the boundary between the northern part of the danish basin and the south-western part of the elevated scandinavian basement is covered by the skagerrak, the sea covering a deep depression (about 500 m deep) known as the norwegian channel (sejrup et al. 1987, 1994, 1998). one of the important discussions concerning the glaciation of denmark during the last stadial focuses on how the ice from norway advanced across the skagerrak about 30 000 years ago. the problem involves the dynamics of the ice stream along the southern coast of norway, the so-called norwegian channel ice stream, and the interaction between the marine and terrestrial parts of the ice cap in southwest norway (larsen et al. 2000). associated problems include the filling of the deep trench in skagerrak, and the termination of marine conditions in skagerrak, vendsyssel, and the northern north sea as well as the kattegat (for locations, see fig. 12). the marine environment referred to as the older yoldia sea, which extended into the vendsyssel region, formed in the late saalian, and the climatic change from a mild climate in the eemian to a glacial 28 ab 58˚n 58˚n 56˚n 56˚n skagen hirtshals frederikshavn mols djursland samsø hven hundested glumsløv ristinge klint bovbjerg jylland jylland læsø anholt langeland 0 100 km50 norway sweden limfjorden skagerrak north sea skagerrak kattegat kattegat vendsyssel mors germanyfig. 13 lodbjerg vendsyssel sd kt msl m sl lim fj or d en i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i denmark 12˚e 12˚e lønstrup lim fjorden jæren n orwegia n tren ch fyn sjælland skåne lønstrup klint 29 climate in the weichselian is recorded in a series of wells drilled in north jylland and the kattegat region (knudsen & lykke-andersen 1982; lykke-andersen 1987; lykke-andersen & knudsen 1991; knudsen 1994). towards the end of the middle weichselian the scandinavian ice sheet over southern norway built up. the ice streams were drained from a main spillway in oslo fjord moving out through the norwegian channel along the coastline of southern norway (larsen et al. 2000). a change in the dynamics of the scandinavian ice sheet over southern norway forced the glaciers to progress south-westward across the norwegian channel. the ice advanced into the northern north sea, where a glacial cover was established about 29 000 years b.p. and lasted until 22 000 years b.p., when the first recurrence of marine conditions (the ‘young yoldia sea’) was recorded (sejrup et al. 1994, 2000). this glacial coverage was probably closely connected with the fall in sea level, amounting to 120 m below present sea level (fairbanks 1989; bard et al . 1993), which could have hampered the active drainage of the norwegian channel ice stream. the ice spread southward over the skagerrak causing the kattegat basin to be dammed by the ice margin and terrestrial areas to be established in the central part of the north sea (sadolin et al. 1997; houmark-nielsen 1999). as a consequence, the kattegat–skagerrak region began to dry up due to the general sea-level fall; this is reflected in the progression from arctic marine conditions in the skærumhede series to brackish and glaciolacustrine environments. this change took place at about 32 000 years b.p. (table 2), and may have been accentuated by the addition of meltwater from the advancing norwegian ice (jessen 1918; sadolin et al. 1997). the dramatic drainage of the lake basin in the kattegat towards the north sea is recorded by a significant erosional unconformity in the sedimentary succession at lønstrup klint (the l/r-unconformity), dated as close to 29 000 years b.p. (sadolin et al. 1997). shortly afterwards, the basin was once again dammed and shallow lacustrine and fluvial environments were established while proglacial thrust faulting was initiated reflecting the relatively fast advance of the ice margin (sadolin et al. 1997). the thin-skinned thrust faulting in the rubjerg knude glaciotectonic complex involved an accretionary wedge extending more than 12 km to the south in front of the advancing ice margin. the lowermost décollement level was situated in the marine clays of the older yoldia sea. after a compression of about 50%, the glaciotectonic complex was formed (pedersen 1987) leaving a large part of the area between lønstrup and hirtshals as a depression corresponding to the ‘hole’ and the rubjerg knude glaciotectonic complex to the ‘hill’, in a ‘hill-andhole’ pair in the sense of aber et al. (1989). subsequently the norwegian ice truncated the glaciotectonic complex and the deposition of the kattegat till formation concealed its structures. the norwegian ice advanced down to a stationary line (figs 1, 12) crossing central denmark from west to east, whose position is inferred from the distribution of the kattegat till formation (houmark-nielsen 1987, 1999, 2003; pedersen & petersen 1997). after its termination at the stationary line (figs 1, 12), the norwegian ice melted back. it was succeeded by the main south-west ice advance of the scandinavian ice sheet, which extended out to the main stationary line (ussing 1903; houmark-nielsen 1987, 2003; pedersen et al. 1988). in northern jylland, the isostatic depression due to the loading of the ice sheet was substantial. the termination of the glaciation in denmark thus resulted in interference between eustatic sea-level rise and isostatic rebound with a complex depositional development during the re-establishment of the younger yoldia sea in the skagerrak– vendsyssel–kattegat region about 17 000 years ago. this may be summarised as a forced regression under progressively falling sea level due to the isostatic rise of the vendsyssel region (richard 1996). the vennebjerg and rubjerg knude hilly islands probably formed part of a larger island archipelago extending out into the north sea. terrestrial conditions were established at the end of the weichselian. at nørre lyngby (fig. 13), a depression was formed above a neotectonic fault zone that predated older dryas time (lykke-andersen 1992). in this depression, lacustrine gyttja and fluvial sand of older dryas and allerød age were deposited; a large number of mammalian remains have been found in these deposits indicating an arctic to sub-arctic reinfacing page: fig. 12. location map. map (a) shows the main part of the danish basin with the surrounding land areas. sdkt is the position of the stationary line for the norwegian ice advance (sdkt is an abbreviation of southern distribution of kattegat till fm). msl is the main stationary line for the scandinavian ice sheet at the glacial maximum in the late weichselian. map (b) gives the position of relevant geographical localities in denmark as well as the location of fig. 13, the geological map of vendsyssel. 30 deer steppe also populated by hunters (jessen & nordmann 1915; aaris-sørensen 1995). during holocene time, the vendsyssel region was affected by isostatic rebound (mertz 1924). at lønstrup klint, this resulted in a 25 m elevation of the heterolithic sediments of the younger yoldia sea. bogs developed in the depressions on the glacial peneplain at the end of the stone age and the beginning of the bronze age (jessen 1918). up to 1.5 m of peat accumulated; when this is exposed in the clif f surface and blocks of peat fall down onto the beach, the peat is locally called martørv (sea-peat). the locality names martørv bakker (sea-peat hill) and moserende (boggully) refer to these deposits. the geomorphology of the cliff is strongly influenced by the thrust-fault structures. the clayey parts of the thrust sheets form ridges that form projections along the coast between gullies that are eroded out in the sandy parts (schou 1949). springs typically well out at the surface between the clayey and sandy units and more incised gullies (render in danish) are formed where the drainage is concentrated. although the location of gullies and the clif f line have retreated about 100 m since a. jessen constructed the first cross-section of lønstrup klint, it has been possible to retain his names in the present cross-section (plate 1). the general erosion rate of the cliff is about 1.3–1.5 m per year (jessen 1918; pedersen 1986b). landslides occur very frequently, particularly at sites where mud diapirs are located in the cliff section. where glaciofluvial deposits dominate the cliff section, there is a marked tendency for aeolian dunes to accumulate on top of the cliff (pedersen 1986b). wind action on the moraine plateau on top of the cliff has eroded the finegrained material away from the till deposits, leaving a stone pavement as the residual trace of the glacially truncated surface. aeolian sand migration intensified about 300–400 years ago (jessen 1918), one of the consequences being the burial and abandonment of the old rubjerg church. the high aeolian dunes on top of rubjerg knude have accumulated during the last 100 years. the rubjerg knude lighthouse was built in 1900 (bendsen 1981) when dunes were less than 10 m high. today the tops of the dunes are close to 100 m above sea level corresponding to a vertical dune accumulation of nearly 50 m. the present-day steep nature of the dunes was probably stimulated by the artificial dune protection fences. however, the steady erosion of the cliff indicates that the lighthouse will fall into the sea about ten years from now. table 2. radiocarbon dates, rubjerg knude and lønstrup klint, vendsyssel, northern denmark stratigraphic unit vendsyssel fm vendsyssel fm vendsyssel fm rubjerg knude fm rubjerg knude fm lønstrup klint fm stortorn fm stortorn fm stortorn fm locality lønstrup klint lønstrup klint lønstrup klint sandrende lønstrup klint sandrende ribjerg mårup kirke stortorn lab. id no. k-858 k-2670 aar-2134 aar-2265 aar-4066 ua-4454 aar-4067 aar-4068 aar-4069 material mollusc mollusc mollusc plant mollusc moss mollusc mollusc mollusc 14c age ka b.p. 13.9 ± 0.2 14.7 ± 0.2 14.5 ± 0.2 30.9 ± 0.5 43.0 ± 1.3 29.2 ± 1.4 29.6 ± 0.4 30.9 ± 0.4 31.3 ± 0.4 calib. age ka b.p.* 16 ± 1 17 ± 1 17 ± 1 33 ± 1 46 ± 3 32 ± 1 33 ± 1 34 ± 1 34 ± 1 13c‰ pdb+ 0.6 –27.3 3.3 –29.1 1.5 1.7 1.3 ref.‡ (1) (2) (3) (4) (5) (4) (5) (5) (5) * calibrated ages are calculated according to bard et al. 1993 and kitagawa & van der plicht (1998). + relative to pdb standard.‡ references: 1: krog & tauber (1974); 2: knudsen (1978); 3: richardt (1996); 4: houmark-nielsen et al. (1996); 5: this study. 31 holocene vendsyssel formation glacigene formations upper cretaceous chalk 20 km5 10 150 10˚30’10˚00’ nørre lyngby mårup løkken rubjerg knude skærumhede skagen frederikshavn hirtshals kattegat skagerrak location of well lø ns tr up k lin t lønstrup sandrende vennebjerg dgu no. 1.287 dgu no. 10.4 dgu no. 8.137 stortorn 57˚30’ dgu no. 8.137 fig. 13. geological map of vendsyssel showing the location of three wells referred to in the text. 32 h ol oc en e m id dl e w ei ch se lia n la te w ei ch se lia n sk æ ru m he de g ro up 11.5 19–15 27–23 30–28 32–30 35–32 fo rm at io n u nc on fo rm ity li th ol og y c la y si lt pe bb le sa nd c ob bl e fo ss ils st ru ct ur es an d gr ai n si ze s d ir ec tio na l el em en ts t hi ck ne ss a ge k a b. p. (c al en da r ye ar s) 30 20 10 0 m recent aeolian dunes main cliff-section unconformity nr. lyngby beds vendsyssel formation mid danish till formation ribjerg formation blå-unconformityeu kattegat till formation rubjerg knude formation eu l/r-unconformity lønstrup klint formation stortorn formation gu gu eu eu clay silty mud dropstones in clay and silty mud sandy mud sand till gravel plant fossil dated shell dated palaeocurrent direction of ice movement erosional unconformity glaciotectonic unconformity f. m. c. eu gu peat and gyttja geological survey of denmark and greenland bulletin 23, 2011, 1–8 1 geological survey of denmark and greenland bulletin 23 • 2011 review of survey activities 2010 edited by ole bennike, adam a. garde and w. stuart watt geological survey of denmark and greenland ministry of climate and energy 22 geological survey of denmark and greenland bulletin 23 keywords geological survey of denmark and greenland, survey organisations, current research, denmark, greenland. cover photographs from left to right 1. small-scale miner with gold concentrate. photograph: peter w.u. appel. 2. work at the microscope. photograph: peter k. warna-moors. 3. many survey employees are engaged in laboratory work. photograph: peter k. warna-moors. 4. development of geological models is becoming increasingly important. photograph: peter k. warna-moors. frontispiece: facing page in 2010 the survey carried out extensive mapping projects in the north sea. the crane is carrying a tow-fish with (1) a side-scan sonar for mapping the seabed and (2) a chirp sonar for mapping the layers below the seabed. photograph: ole bennike. chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, department of earth sciences, university of aarhus; minik rosing, geological museum, university of copenhagen; finn surlyk, department of geography and geology, university of copenhagen scientific editors: ole bennike, adam a. garde and w. stuart watt editorial secretaries: jane holst and esben w. glendal referees: (dk = denmark etc.; numbers refer to first page of reviewed article): anonymous (21, 37, 41, 53, 73), niels balling, dk (49); jason box, usa (73); michele crosetto, e (41); gregers dam, dk (61); david lundbek egholm, dk (69); synnøve elvevold, n (57); ida fabricius, dk (13); rasmus fensholt, dk (81); tom frisch, cnd (69); svend funder, dk (29); rikke harlou, dk (57); jens havskov, n (49); claus heilmann-clausen, dk (61); rasmus jakobsen, dk (45); john a. korstgård, dk (53, 77); gunnar larsen, dk (45); nicolaj krog larsen, s (33); kaj lax, s (77); ole bjørslev nielsen, dk (17); bent odgaard, dk (29); odleiv olesen, n (81); asger ken pedersen, dk (65); gunver krarup pedersen, dk (17); sandra piazolo, s (65); peter sandersen, dk (25); ulf sivhed, s (9); inga sørensen, dk (21); jette sørensen, dk (25); svend stouge, dk (9); szymon uścinowicz, pl (37); ole v. vejbæk, dk (13); jacob clement yde, n (33). illustrations: stefan sølberg, with contributions from jette halskov, eva melskens and benny m. schark layout and graphic production: annabeth andersen printers: rosendahls . schultz grafisk a/s, albertslund, denmark manuscripts received: 21 december 2010 – 6 may 2011 final versions approved: january–may 2011 printed: 15 july 2011 issn 1603-9769 (review of survey activities) issn 1604-8156 (geological survey of denmark and greenland bulletin) isbn 978-87-7871-313-1 citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 23, 84 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk and at www.geus.dk/publications/bull © de nationale geologiske undersøgelser for danmark og grønland (geus), 2011 for the full text of the geus copyright clause, please refer to www.geus.dk/publications/bull 3 44 ghana uganda tanzania mozambique greenland united kingdom norway faroe islands kenya spain sweden denmark bahrain belgium angola democratic republic of the congo cameroun ethiopia nigeria bolivia brazil togo mauritius malawi namibia botswana lesotho south africa zimbabwe zambia yemen swaziland seychelles germany the netherlands 7 review of survey activities 2010 f.g. christiansen 9 shale gas investigations in denmark: lower palaozoic shales on bornholm n.h. schovsbo, a.t. nielsen, k. klitten, a. mathiesen and p. rasmussen 13 mapping porosity anomalies in deep jurassic sandstones – an example from the svane-1a area, danish central graben t. abramovitz 17 differentiation of palaeogene sand by glauconitic and geochemical fingerprinting, siri canyon, danish north sea m. olivarius, c. knudsen and j.b. svendsen 21 geological characterisation of potential disposal areas for radioactive waste from risø, denmark p. gravesen, m. binderup, b. nilsson and s.a.s. pedersen 25 a digital, spatial, geological model of the miocene in jylland, denmark m. kristensen, t. vangkilde-pedersen and e.s. rasmussen 29 a new middle pleistocene interglacial sequence from måløv, sjælland, denmark o. bennike, e. lindgård, h.j. granat, r.c. preece and f. viehberg 33 mapping of raw materials and habitats in the danish sector of the north sea j.b. jensen, s. borre, j.o. leth, z. al-hamdani and l.g. addington 37 postglacial, relative shore-level changes in lillebælt, denmark o. bennike and j.b. jensen 41 detection of terrain changes in southern denmark using persistent scatterer interferometry s.a.s. pedersen, g. cooksley, m. gaset and p.r. jakobsen 45 does road salt affect groundwater in denmark? s.m. kristiansen, f.d. christensen and b. hansen 5 vietnam phillipines india 49 comprehensive nuclear-test-ban treaty – a peace-keeping initiative with scientific impact t.b. larsen, p.h. voss, t. dahl-jensen and s. gregersen 53 free, online danish shallow geological data m. hansen and b. pjetursson 57 remnants of mesoarchaean oceanic crust in the tartoq group, south-west greenland k. szilas, v.j. van hinsberg, a.f.m. kisters, t.f. kokfelt, a. scherstén and b.f. windley 61 palaeogene deposits in north-east greenland h. nøhr-hansen, l.h. nielsen, e. sheldon, j. hovikoski and p. alsen 65 analysis of palaeogene strike-slip tectonics along the southern east greenland margin (sødalen area) p. guarnieri 69 kennedy channel and its geophysical lineaments: new evidence that the wegener fault is a myth t.m. rasmussen and p.r. dawes 73 programme for monitoring of the greenland ice sheet (promice): first temperature and ablation records d. van as, r.s. fausto and the promice project team 77 dodex – geoscience documents and data for exploration in greenland p. riisager, m. pedersen, m.s. jørgensen, f. schjøth and l. thorning 81 quality control of airborne geophysical data from the eu mining sector support programme, ghana t.m. rasmussen, l. thorning, a.v. olesen and f. schjøth geus working areas 2010. orange areas are covered in this volume. for further information on other working areas please refer to our website: www.geus.dk/international 66 7 2010 was a good and stable year for the geological survey of denmark and greenland (geus) with focus on research, often in international collaboration. despite the continued effects of the international financial crisis, which has had serious implications for many of our national and international partners, geus has had a period with many new projects and successful completion of many projects. this is also reflected in the present eighth annual issue of review of survey activities which describes selected projects that geus and its partners carry out in denmark, greenland and internationally. together with the previous seven published issues, it provides a good overview of the survey’s range of research and advisory activities. it contains a total of 19 four-page papers: 12 on denmark, six on greenland, and one project in ghana. energy policy is again high on the political agenda in denmark. the government presented a new energy 2050 strategy with strong emphasis on the reduction of co2 emission. the strategy depends on a stable supply of and income from oil and gas in the north sea during a long transition period before most of denmark’s energy supply becomes co2 neutral, competitive and stable. geus’ research lies within a variety of different aspects of energy as well as of climate development, climate monitoring and adaptation to climate changes. three papers concentrate on various aspects of petroleum geology in denmark. one of them provides an overview of a core-drilling project in lower palaeozoic shales on bornholm as an unconventional shale gas resource analogue. another paper describes mapping of very deep jurassic targets in the svane-1 area in the north sea, and a third paper gives a geochemical fingerprinting of palaeogene reservoir sands from the siri canyon in the north sea and discusses the implications from being able to distinguish between in situ and mobilised sand. geus works on many other aspects of the geology of denmark, such as groundwater, climate and the environment including issues where geology is important to society. seven papers ranging from applied geology to more basic research are found in this volume. decisions on disposal of danish lowand intermediate-level radioactive waste have to be taken in the coming years, and a number of key geological parameters are used for the final selection of a permanent depository. the work has resulted in the selection of 22 areas, of which six are preferred. the miocene succession in jylland contains several large groundwater bodies, and a 3d model is important for future planning; this is described in another paper. with a continued need for raw materials to large infrastructure projects, systematic mapping and understanding of available marine resources are important. the results from a project in the north sea carried out for the danish nature agency are presented in one paper. two papers describe basic research on a new pleistocene interglacial sequence from sjælland and on postglacial relative shore-level changes in lillebælt. another paper presents results of the detection of terrain changes using satellite data with south-western jylland as a case. one paper discusses how sensitive the groundwater quality in denmark is to the use of road salt in winter. the danish contribution to the nuclear-test-ban treaty is also described, demonstrating how useful the data from the monitoring systems are for understanding earthquakes. in 2010, there was a high level of field activities in greenland. in addition to major projects in southern west greenland, south-east greenland and north-east greenland there were many smaller activities in other areas. the work in southern west greenland continued, and results are described in a paper on the tartoq group, a possible very old slab of oceanic crust. field work and shallow core drilling in north-east greenland continued in 2010. in this issue results on previously almost unknown palaeogene sand are presented, including new critical information on age and depositional environment. structural data from the sødalen area in southern east greenland are presented in a paper that concentrates on unravelling the strike-slip tectonics in palaeogene time. one paper adds additional evidence from magnetic data contrareview of survey activities 2010 flemming g. christiansen deputy director © geus, 2010. geological survey of denmark and greenland bulletin 23, 7–8. open access: www.geus.dk/publications/bull 88 dicting the existence of a major structural feature (the wegener fault) in the kennedy channel between north greenland and ellesmere island in canada. studies of the ice sheet and glaciers in greenland have attracted international interest over many years due to the possible implications of a rising sea level. geus is involved in many glaciological and meteorological projects and monitoring programmes. one paper gives a presentation on the large-scale programme for monitoring of the greenland ice sheet (promice) with description of the weather station network and preliminary temperature data. easy access to comprehensive and updated information and data is a very important part of the work geus carries out in denmark and greenland. this is the topic of two papers, one on free, online danish geological data where the jupiter database currently includes information from more than 260 000 shallow wells. the other paper gives a description of dodex (geoscience documents and data for exploration in greenland), which is an interactive web application, which gives the public and mining companies easy access to all non-confidential reports relevant to mineral exploration. internationally geus works in many different countries with many project types. the last paper in this issue is about work in ghana where geus has been active for many years with capacity building and geological and geophysical projects. the paper gives an overview of quality control of airborne geophysical data. geological survey of denmark and greenland bulletin 6, 73-88 73geological survey of denmark and greenland bulletin 5, 73–88 © geus, 2004 jurassic dinoflagellate cyst stratigraphy of hold with hope, north-east greenland stefan piasecki, michael larsen, jens therkelsen and henrik vosgerau dinoflagellate cysts of the middle–upper jurassic succession on northern hold with hope have been studied in order to establish a biostratigraphic framework and to date the succession. the pelion formation is characterised by abundant chytroeisphaeridia hyalina and sentusidinium spp., with some ctenidodinium thulium and paragonyaulacysta retiphragmata in the lower part. mendicodinium groenlandicum appears higher in the formation followed by trichodinium scarburghense in the upper part. the succeeding payer dal formation contains scriniodinium crystallinum, rigaudella aemula and leptodinium subtile in the lower part and dingodinium jurassicum and prolixosphaeridium granulosum in the uppermost part. the bernbjerg formation contains abundant sirmiodinium grossii and gonyaulacysta jurassica. adnatospahaeridium sp., cribroperidinium granuligerum, glossodinium cf. dimorphum and scriniodinium irregulare appear in the lower part of the formation, followed by avellodinium spp. in the highest part. the dinoflagellate cyst assemblages in the pelion formation indicate an early–late callovian age (c. apertum – p. athleta chronozones). this is supported by ammonites in the lower part of the formation, which refer to the c. apertum and p. koenigi chronozones. a significant hiatus, from late callovian to middle oxfordian, is present between the pelion formation and the overlying payer dal formation. the age of the payer dal formation is middle oxfordian to earliest late oxfordian (c. tenuiserratum – a. glosense chronozones). the payer dal formation is conformably overlain by the bernbjerg formation of late oxfordian to possibly earliest kimmeridgian age (a. glosense – p. baylei chronozones). the a. glosense chronozone is also documented by abundant ammonites in the lowermost part of the formation. keywords: ammonites, dinoflagellate cysts, jurassic, north-east greenland, stratigraphy s.p., m.l., j.t.* & h.v.‡, geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: sp@geus.dk present addresses: *skude & jacobsen, næstvedvej 1, dk-4760 vordingborg, denmark. ‡roskilde amt, køgevej 80, dk-4000 roskilde, denmark. the recognition of middle–upper jurassic sediments on northern hold with hope added a missing link to the chain of jurassic sedimentary exposures along the east coast of greenland (figs 1, 2; stemmerik et al. 1997; kelly et al. 1998; larsen et al. 1997; vosgerau et al. 2004, this volume). sedimentological and biostratigraphical analysis of the succession formed the basis for correlation with lithostratigraphical units in wollaston forland and jameson land, and subdivision into the pelion, payer dal and bernbjerg formations (fig. 3). a new member of the pelion formation, the spath plateau member, was erected (vosgerau et al. 2004, this volume). correlation was based on very few, poorly preserved middle jurassic ammonites in situ in the lower sandstone-dominated part of the succession, and more abundant upper jurassic ammonites of the upper oxfordian, the a. glosense zone, in the mudstonedominated upper part of the succession. the content of dinoflagellate cysts was studied in order to improve the biostratigraphic dating of the succession, and to geus bulletin no 5.pmd 29-10-2004, 11:1473 74 improve the knowledge of jurassic dinoflagellate cysts in this region in general. the results reported here allow correlation with corresponding assemblages from store koldewey and hochstetter forland in the north and jameson land – milne land in the south (fig. 1). geological setting the late palaeozoic – mesozoic extensional basin complex in east greenland is approximately 700 km long in a north–south direction. jurassic sediments are present and exposed from jameson land in the south to store koldewey in the north (surlyk 1977). in the northern part of the rift system, e.g. the wollaston forland basin, rifting was initiated in middle jurassic time, and marine bajocian–bathonian sandstones onlap caledonian basement rocks or permian carbonates (vischer 1943; surlyk 1978). deposition took place on the hangingwall of w–sw-tilted fault blocks. jurassic rifting culminated in the volgian with strong rotational block faulting (surlyk 1978). during this episode the wide original fault blocks, defining the wollaston forland basin, were divided into smaller blocks (vischer 1943; surlyk 1978). a similar tectonic development may have occurred in the geographical society ø and traill ø area towards the south (donovan 1957; price & whitham 1997). the cretaceous period was generally characterised by subsidence controlled by thermal contraction (surlyk et al. 1981; price & whitham 1997). the east greenland rift basin complex was uplifted during the cenozoic. sediments of jurassic age were first recognised on hold with hope by stemmerik et al. (1997). they are limited to the north coast of hold with hope from stensiö plateau to steensby bjerg (fig. 2), where they occur on the hangingwall of small fault blocks that dip mainly to the west and south-west. bedding planes within the triassic and jurassic seem to be parallel, whereas the boundary with the overlying cretaceous succession is an angular unconformity (vosgerau et al. 2004, this volume). the thickness of the jurassic succession varies significantly depending on its position on the hangingwall and the depth of cretaceous erosion. the jurassic succession includes shallow marine sandstones of the pelion and payer dal formations (vardekløft group), and offshore transition – lower shoreface heteroliths and offshore mudstones of the bernbjerg formation, hall bredning group (fig. 3). the spath plateau member of the pelion formation was erected to accommodate sandy heteroliths and fig. 1. locality map of east greenland and eastern north greenland. the white region illustrates the permanent inland ice cap of greenland, the grey areas are ice-free. hold with hope is located between 73ºn and 74ºn. scoresby sund st. koldewey kuhn ø fig. 2 wollaston forland hold with hope geographical society ø traill ø milne land jameson land hochstetter forland 200 km 20°w28°w36°w42°w 12°w 70°n 28°w 20°w 72°n 74°n 76°n 78°n 80°n 12°w 82°n geus bulletin no 5.pmd 29-10-2004, 11:1474 75 offshore mudstones that contrast with the generally coarse-grained sandstone facies of the pelion formation (vosgerau et al. 2004, this volume). the succession on hold with hope resembles the well-known jurassic succession in the wollaston forland and jameson land basins towards the north and south. the middle jurassic pelion formation is c. 190 m thick, the upper jurassic payer dal formation is 50–80 m thick and the bernbjerg formation is estimated to be c. 130 m thick (fig. 4; vosgerau et al. 2004, this volume). samples and methods the dinoflagellate cysts have been studied in three sections (fig. 4), in combination with a number of geographically and stratigraphically scattered samples on northern hold with hope. the main area of exposure is located on the northern and western slopes of steensby bjerg towards gael hamke bugt and along the gulelv river (fig. 2, locality 1). samples from a number of short, vertical sections are combined into a composite section representing the entire succession (fig. 2, locality 1, sections a–e). the payer dal formation was also sampled at the sortelv river, south of steensby bjerg (fig. 2, locality 2). samples from a third section through the pelion formation at stensiö plateau (fig. 2, locality 3) provide good supplementary material from the lowermost part of the succession, which is poorly represented in the section at steensby bjerg. most of the analysed samples are from finegrained thin beds or lamina in the otherwise coarsegrained, sandy pelion and payer dal formations. the number of samples and their stratigraphical distribufig. 2. geological map of the northern hold with hope which illustrates the distribution of the studied jurassic succession. localities 1, 2 and 3 are marked; the succession at locality 1 has been compiled from a number of short laterally correlated sections (1a–1e). ice undifferentiated superficial deposits dolerite sill plateau basalt normal fault inferred fault locality paleocene lower cretaceous middle and upper jurassic lower triassic permian 2 km crystalline basement 1b 1a 1c 1d 1e 2 3 ■■ ■■ 3 ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ steensby bjerg gael hamke bugt diener bjerg sortelv spath plateau stensiö plateau 21°00' 21°15'w 74°00'n 73°55'n hold with hope g ul el v ■ ■ ■■ ■■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ geus bulletin no 5.pmd 29-10-2004, 11:1475 76 tion are controlled by the occurrence and accessibility of these fine-grained beds. in contrast, the shale of the bernbjerg formation provides productive samples throughout the formation. standard palynological preparation has been performed on most samples. a minority of the samples were prepared by the tank-preparation method (poulsen et al. 1990). both methods involve treatment with hydrofluoric (hf) and hydrochloric acids (hcl) followed by filtering at 20 µm mesh size, short oxidation by nitric acid (hno 3 ) and washing in low concentration potassium hydroxide (koh). biostratigraphy the ammonites and dinoflagellate cysts have been analysed and correlated to the boreal ammonite and dinoflagellate stratigraphy, i.e. east greenland stratigraphy (callomon 1993; milner & piasecki 1996; piasecki 1996; piasecki & stemmerik 2004, this volume; piasecki et al. 2004, this volume). ammonites ammonites are very restricted in the jurassic succession on hold with hope, and only three horizons have been dated and correlated with the standard boreal ammonite stratigraphy (callomon 1993). a specimen referred to cadoceras cf. breve (j.h. callomon and p. alsen, personal communications 1997) was found 10 m a. mutabilis r. cymodoce p. baylei bernbjerg formation payer dal formation sandstone lower sandstone unit spath plateau member a. rosenkrantzi a. regulare a. glosense c. tenuiserratum c. densiplicatum c. cordatum q. mariae q. lamberti p. athleta e. coronatum k. jason p. koenigi c. nordenskjoeldi c. apertum c al lo vi an o xf or di an k im m er id gi an u pp er ju ra ss ic m id dl e ju ra ss ic l l m pelion formation m u u s. calloviense a. serratum chronozones lithology lithostratigraphy heterolithic sandstone ammonite dinoflagellate cyst mudstone fig. 3. schematic correlation of the jurassic succession on northern hold with hope. the lithostratigraphical units are correlated with the middle to upper jurassic chronozonation on the basis of ammonites and dinoflagellate cysts. points of correlation to chronozones are indicated by ammonite or dinoflagellate cyst signatures. the subdivision of the chronozones corresponds to the ammonite faunas in the biozonation. facing page: fig. 4. simplified sedimentological logs of the jurassic succession from localities 1, 2 and 3 on northern hold with hope. the formal and informal lithostratigraphic units are indicated together with the ammonite horizons; l.s., lower shale. geus bulletin no 5.pmd 29-10-2004, 11:1476 77 silt sand pebbles f m c locality 1 locality 2 lithology locality 3 poor exposure 427851 427850 427859 427810 427809 427808 427796 427793 427790 427780 427840 444854 444832 444855 444857 433158 433157 433859 ggu sample no. 433159 433858 444856 433868 444830 433869 433865 427833 427836 427730 427717 427729 427729 427818 427782 427705 427858 ggu sample no. siltstone poor exposure a. glosense a. glosense poor exposure up pe r sa nd st on e un it l. s. un it sp at h pl at ea u m em be r pe lio n fo rm at io n pe lio n fo rm at io n c . a pe rt um – p . k oe ni gi be rn bj er g fo rm at io n pa ye r d al f or m at io n lo w er sa nd st on e un it c . a pe rt um – p . k oe ni gi k. ja so n p. at hl et a a. g lo se ns e c . t en ui se rr at um p. ba yle i 50 100 150 200 250 300 350 m 0 silt sand pebbles f m c 0 30 m p. koenigi spath plateau member c. apertum silt sand pebbles f m c 0 50 m silty sandstone sandstone pebbly sandstone pebble lag structures horizontal lamination planar bedding wave ripple planar cross-bedding trough cross-bedding ammonite/zone hummocky cross-stratification cross-lamination lo w er sa nd st on e un it pa ye r d al f or m at io n c . t en ui se rr at um – a . g lo se ns e geus bulletin no 5.pmd 29-10-2004, 11:1477 78 above the base of the lower sandstone unit in the pelion formation and indicates the cadoceras apertum zone (fig. 4). a poorly preserved ammonite referred tentatively to cadoceras septentrionale (p. alsen, personal communication 1998) in the lowermost spath plateau member of the pelion formation indicates the proplanulites koenigi zone. much higher in the succession, in the basal bernbjerg formation, the presence of amoeboceras ilovaiskii (j.h. callomon and p. alsen, personal communications 1997) indicates the amoeboceras glosense zone. these three ammonite horizons occur at separate localities (fig. 4). the c. apertum zone is identified in the succession at stensiö plateau (locality 3), the p. koenigi zone is identified in the succession at gulelv (locality 1) and the a. glosense zone is identified in the section at sortelv (locality 2). a calcareous concretion with a specimen of cranocephalites sp. (c. pompeckji zone) is reworked into the cretaceous basal conglomerate. the ammonite data thus indicate that parts of the lower pelion formation are equivalent to the c. apertum – p. koenigi chronozones, lower callovian, and parts of the lower bernbjerg formation are equivalent to the a. glosense chronozone, upper oxfordian. a more detailed stratigraphical framework is provided by the more consistently occurring dinoflagellate cysts. dinoflagellate cysts the dinoflagellate cyst data are described below in relation to five lithostratigraphic units, as presented by vosgerau et al. (2004, this volume). pelion formation, lower sandstone unit the dinoflagellate cyst assemblages are of low to moderate diversity and density in the samples from this coarse-grained unit. the most diverse assemblages were recovered from the succession at stensiö plateau (figs 2, 4, 5, locality 3). many of the species in this assemblage are known from strata in east greenland older than the early callovian age indicated here by ammonites (milner & piasecki 1996). three assemblages have been distinguished in this unit, based on a limited number of samples. a lower assemblage of poor diversity with frequent chytroeisphaeridia hyalina and sentusidinium sp. d (fensome 1979) is followed by a middle assemblage of higher diversity with abundant sirmiodinium grossii, valensiella dictydia and sentusidinium spp. the third and uppermost assemblage, above the ammonite horizon of the c. apertum zone, is moderately to highly diverse and contains abundant chytroeisphaeridia hyalina, rhynchodiniopsis cladophora, r. cf. cladophora and pareodinia pachyceras (fig. 5). the corresponding succession at steensby bjerg (figs 2, 6, locality 1) contains a very poor dinoflagellate cyst assemblage and chytroeisphaeridia hyalina is the only frequent species. however, also at this locality slightly more species appear in the uppermost strata of the unit, thus showing an upwards increase in diversity. correlation. the succession of species appearances up through the lower sandstone unit of the pelion formation in the stensiö plateau succession (locality 3) does not yield any significant stratigraphic information. however, the abundance of chytroeisphaeridia hyalina combined with the earliest appearance of fromea tornatilis, pareodinia prolongata, aldorfia aldorfensisandkallosphaeridium sp. in this unit are considered indicative of the c. apertum chronozone based on comparison to the dinoflagellate records in jameson land and store koldewey (milner & piasecki 1996; piasecki et al. 2004, this volume). this is also in accordance with the ammonite record in this succession. the poor dinoflagellate assemblage from the ‘lower sandstone unit’ at steensby bjerg (locality 1) does not provide clear correlation but contains some characteristic species, e.g. paraevansia brachythelis which has its lowest record in the c. apertum chronozone on store koldewey (piasecki et al. 2004, this volume). several species that are restricted to the upper assemblage of the stensiö plateau succession are also limited to the topmost strata of the corresponding unit in the succession at steensby bjerg: aldorfia aldorfensis, lithodinia planoseptata, ctenidodinium thulium and pareodinia prolongata. however, other species from the upper assemblage in the stensiö plateau succession (c. apertum chronozone at locality 3) appear for the first time at a stratigraphically higher level in the steensby bjerg succession (locality 1). this may reflect the restricted material and data from this unit in the steensby bjerg succession (locality 1). age. the age of the ‘lower sandstone unit’ of the pelion formation is early callovian, equivalent to the c. apertum – p. koenigi chronozones based on ammonites and dinoflagellate cysts. geus bulletin no 5.pmd 29-10-2004, 11:1478 79 depositional environment. the presence of a low diverse assemblage with limbicysta bjaerkei in the basal strata combined with significant, upwards increasing diversity indicate that deposition of this unit began in a marginal marine environment and changed to deposition in a fully marine environment. the preferred habitat of l. bjaerkei is non-marine (bailey & hogg 1995) but it also has been recorded in restricted marine dinoflagellate cyst assemblages, for example in the basal strata of the payer dal formation in hochstetter forland (piasecki & stemmerik 2004, this volume). here, l. bjaerkei occurs together with the marine fauna immediately above non-marine–brackish sediments. pelion formation, spath plateau member, lower shale unit the diversity and especially abundance of dinoflagellate cysts reach a maximum in the basal mudstone of the spath plateau member. in the stensiö plateau succession (locality 3), the composition of the assemblage is not significantly different from the highest assemblage in the unit below. however, in the steensby bjerg succession (locality 1), several species appear stratigraphically delayed compared to the stensiö plateau succession and their appearance in this ‘lower shale unit’ produces a local, significant increase in the diversity (fig. 6). chytroeispharidia hyalina is very abundant at both localities together with frequent sirmiodinium grossii, sentusidinium pelionense, rhynchodiniopsis cladophora, r. cf. cladophora and sentusidinium sp. d (fensome 1979). correlation. the ammonite biostratigraphy shows that the mudstone is within or above the c. apertum and the p. koenigi chronozones at localities 1 and 3, respectively. the dinoflagellate biostratigraphy suggests that this mudstone is of the same age at localities 1 and 3, i.e. equivalent to the p. koenigi chronozone, but the stratigraphic resolution does not exclude the possibility that the basal mudstone at locality 3 may include strata from the c. apertum chronozone. this is the stratigraphical lower limit based on ammonites (fig. 4). it is possible that the mudstone is diachronous. the ammonite found in sandstone at the lithostratigraphic transition to the basal mudstone of the spath plateau member at locality 1 (fig. 4), is referred to the proplanulites koenigi zone. most of the dinoflagellate species that appear just above the ammonite at this locality, are reported to appear for the first time in or near the c. apertum chronozone. the highest occurrence of paragonyaulacysta retiphragmata is found at the same level in both successions (localities 1, 3) and indicates the p. koenigi chronozone based on its last occurrence in the jameson land basin (milner & piasecki 1996). the highest occurrence of kallosphaeridium hypornatum in jameson land is also in the p. koenigi chronozone. pareodinia stegasta appears in the basal mudstone as it does in a stratigraphically comparable transgressive shale unit on store koldewey (piasecki et al. 2004, this volume). the lower boundary of the spath plateau member is a major drowning surface overlain by mudstone both on hold with hope and on store koldewey (piacecki et al. 2004, this volume; vosgerau et al. 2004, this volume). age. the age of the ‘basal shale unit’ of the spath plateau member is early callovian, equivalent to the c. apertum – p. koenigi chronozones. depositional environment. the maximum diversity and density of dinoflagellate cysts in the middle jurassic succession occur in this unit and indicate deposition of shelf mudstone in a fully marine environment during flooding. pelion formation, spath plateau member, upper sandstone unit samples are available only from the succession at steensby bjerg (locality 1). the dinoflagellate assemblage is moderately rich and diverse. the bulk of the species are the same as in the shale below, but are combined with more species higher in the succession that typically appear in the callovian. chytroeisphaeridia hyalina, gonyaulacysta jurassica, rhynchodiniopsis cladophora and sentusidinium spp. are most frequent. mendicodinium groenlandicum appears in the lower part of the unit and tubotuberella eisenackii and trichodinium scarburghense appear higher in the unit. correlation. the overall callovian dinoflagellate assemblage provides few stratigraphical markers. the unit is stratigraphically restricted downwards by the presence of lower callovian dinoflagellate cysts and ammonites (p. koenigi chronozone) in the shale unit below. records from the jameson land basin indicate that mendicodinium groenlandicum appears in the k. geus bulletin no 5.pmd 29-10-2004, 11:1479 80 h old w ith h ope, locality 3, stensiö plateau metres 5030 sample height age stage formation 45.00 37.00 31.25 28.00 23.00 21.00 7.50 6.50 2.25 1.00 444853 444857 444856 444855 433869 444832 444830 433868 444854 433865 cretaceous steensby bjerg middle jurassic callovian pelion 1 tasmanites spp. 2 valensiella dictydia 3 solisphaeridium ankyleton 4 leiofusa jurassica 5 rhynchodiniopsis cf. regalis? 6 sentusidinium spp. 7 pareodinia halosa 8 chytroeisphaeridia hyalina 9 valensiella ovula 10 sentusidinium cf. pelionense 11 sentusidinium sp. d (fensome 1979) 12 fromea tornatilis 13 atopodinium spp. 14 valvaeodinium hanneae 15 ctenidodinium thulium 16 nannoceratopsis plegas var. dictyornata 17 paraevansia spp. 18 sirmiodinium grossii 19 lithodinia spongiosa 20 cyclopsiella spp. 21 rhynchodiniopsis cladophora 22 tubotuberella spp. 23 ambonosphaera calloviana 24 lithodinia spp. 25 valvaeodinium leneae 26 sentusidinium pelionense 27 paragonyaulacysta retiphragmata 28 pareodinia spp. 29 pterospermopsis sp. a (fensome 1979) 30 valensiella spp. 0 2010 40 ggu sample no. g e u s b ulletin no 5.pm d 29-10-2004, 11:14 80 81 fig. 5. d istrib u tio n ch art o f d in o flagellate cysts in th e ju rassic su ccessio n at lo cality 3, sten siö p lateau , n o rth ern h o ld w ith h o p e. t h e first ap p earan ces o f sp ecies are stratigrap h ically arran ged . t h e ju rassic su ccessio n is o verlain u n co n fo rm ab ly b y c retaceo u s strata o f th e steen sb y b jerg fo rm atio n (k elly et a l. 1998) – see sam p le at 45 m . 31 solisphaeridium spp. 32 lithodinia planoseptata 33 aldorfia aldorfensis 34 atopodinium haromense 35 kallosphaeridium hypornatum 36 gonyaulacysta jurassica 37 pareodinia "granulata" 38 pareodinia prolongata 39 endoscrinium galeritum 40 pareodinia pachyceras 41 kallosphaeridium praussii. 42 pilosidinium fensomei 43 rhynchodiniopsis cf. cladophora 44 chlamydophorella ectotabulata 45 gonyaulacysta cf. helicoidea 46 micrhystridium spp. 47 sentusidinium sparsibarbatum 48 mendicodinium spp. 49 pareodinia stegasta 50 paraevansia brachythelis 51 lithodinia cf. callomonii 52 gonyaulacysta pectinigera 53 escharisphaeridia rudis 54 veryhachium sortehatense 55 micrhystridium cf. deflandrei 56 tubotuberella eisenackii 57 solisphaeridium cf. stimuliferum 58 paragonyaulacysta sp. (fensome 1979) 59 batioladinium pelliferum 60 escharisphaeridia spp. 61 hystrichodinium spp. 62 chytroeisphaeridia spp. a lph a bet ic a l spec ies list 33 a ldorfia aldorfensis 23 a m bonosphaera calloviana 34 a topodinium harom ense 13 a topodinium spp. 59 b atioladinium pelliferum 44 c hlam ydophorella ectotabulata 8 c hytroeisphaeridia hyalina 62 spp. c hytroeisphaeridia 15 c tenidodinium thulium 20 spp. c yclopsiella 39 e ndoscrinium galeritum 60 spp. e scharisphaeridia 53 e scharisphaeridia rudis 12 f rom ea tornatilis 45 cf. g onyaulacysta helicoidea 36 g onyaulacysta jurassica 52 g onyaulacysta pectinigera 61 h ystrichodinium spp. 35 k allosphaeridium hypornatum 41 m k allosphaeridiu praussii 4 leiofusa jurassica 51 lithodinia cf.callom onii 32 lithodinia planoseptata 19 lithodinia spongiosa 24 lithodinia spp. 48 m endicodinium spp. 55 m icrhystridium cf. deflandrei 46 m icrhystridium spp. 16 n annoceratopsis plegas var. dictyornata 29 p terosperm opsis sp. a (f ensom e 1979) 50 p araevansia brachythelis 17 p araevansia spp. 27 p aragonyaulacysta retiphragm ata 5837 p areodinia "granulata" 7 p areodinia halosa 40 p areodinia pachyceras 38 p areodinia prolongata 28 p areodinia spp. 49 p areodinia stegasta 42 p ilosidinium fensom ei 43 r hynchodiniopsis cf.cladophora 5 r hynchodiniopsis cf. regalis ? 21 r hynchodiniopsis cladophora 10 cf. s entusidinium pelionense 26 s entusidinium pelionense 11 s entusidinium sp. d (f ensom e 1979) 47 s entusidinium sparsibarbatum 6 s entusidinium spp. 18 s irm iodinium grossii 3 s olisphaeridium ankyleton 57 s olisphaeridium cf. stim uliferum 31 s olisphaeridium spp. 1 tasm anites spp. 56 tubotuberella eisenackii 22 tubotuberella spp. 2 v alensiella dictydia 9 v alensiella ovula 30 v alensiella spp. 14 v alvaeodinium hanneae 25 v alvaeodinium leneae 54 v eryhachium sortehatense p aragonyaulacysta sp. (f ensom e 1979) > 50 specim ens 2–50 specim ens 2–4 specim ens 1 specim en 5–19 specim ens g e u s b ulletin no 5.pm d 29-10-2004, 11:14 81 82 hold with hope, locality 1, steensby bjerg m et re s 30 0 20 0 10 0 0 sa m pl e he ig ht a ge st ag e fo rm at io n 360.00 349.00 262.00 248.00 221.00 219.00 202.00 168.00 134.00 121.00 77.00 70.00 60.00 51.00 39.00 36.00 33.00 30.00 27.00 10.00 9.00 g g u s am pl e no . 427851 427850 427859 427858 427810 427809 427808 427796 427793 427790 427780 427840 427833 427836 427730 427729 427818 427782 427705 427717 427729 la te ju ra ss ic m id dl e ju ra ss ic k im m .? m id dl e – u pp er o xf or di an c al lo vi an be rn bj er g pa ye r d al pe lio n pe lio n – sp at h pl at ea u m em be r 1 li m bi cy st a bj ae rk ei 2 p ar ae va ns ia b ra ch yt he lis 3 c hy tr oe is ph ae rid ia cf . c er as te s 4 r hy nc ho di ni op si s cf . c la do ph or a 5 c hy tr oe is ph ae rid ia h ya lin a 6 e sc ha ris ph ae rid ia cf .p oc oc ki i 7 v al en si el la d ic ty di a 8 e sc ha ris pa ha er ia la ev ig at a 9 a to po di ni um sp p. 10 s en tu si di ni um c f. pe lio ne ns e 11 s irm io di ni um g ro ss ii 12 p ar eo di ni a "g ra nu la ta " 13 n an no ce ra to ps is p el lu ci da 14 v al en si el la o vu la 15 g on ya ul ac ys ta c f. he lic oi de a 16 p ar eo di ni a pr ol on ga ta 17 c te ni do di ni um th ul iu m 18 li th od in ia sp p. 19 a ld or fia a ld or fe ns is 20 li th od in ia p la no se pt at a 21 a to po di ni um p ol yg on al e 22 k al lo sp ha er id iu m h yp or na tu m 23 b ar ba ta cy st a cr eb er ba rb at a 24 a lg ae in de t. 25 g on ya ul ac ys ta ju ra ss ic a 26 r hy nc ho di ni op si s cl ad op ho ra 27 tu bo tu be re lla s pp . 28 p ar ag on ya ul ac ys ta r et ip hr ag m at a 29 v al en si el la s pp . 30 li th od in ia s po ng io sa 31 p ar eo di ni a ha lo sa 32 k al lo sp ha er id iu m p ra us si i 33 c yc lo ps ie lla sp p. 34 s en tu si di ni um sp . d ( f en so m e 19 79 ) 35 s ur cu lo sp ha er id iu m sp p. 36 s en tu si di ni um sp p. 37 m en di co di ni um g ro en la nd ic um 38 p ar eo di ni a ce ra to ph or a 39 v er yh ac hi um s pp . 40 tu bo tu be re lla d an ge ar di i 41 li th od in ia cf . s po ng io sa 42 v al en si el la s p. ( f en so m e 19 79 ) geus bulletin no 5.pmd 29-10-2004, 11:1482 83 43 fromea tornalis 44 lithodinia jurassica 45 escharisphaeridia rudis 46 kallosphaeridium spp. 47 durotrigia spp. 48 nannoceratopsis plegas var. dictyornata 49 lithodinia valensi 50 mendicodinium "granulatum" 51 pareodinia pachyceras 52 chlamydophorella ectotabulata 53 atopodinium haromense 54 ambonosphaera calloviense 55 sentusidinium pelionense 56 gonyaulacysta spp. 57 pareodinia stegasta 58 tubotuberella eisenackii 59 trichodinium scarburghense 60 gonyaulacysta eisenackii 61 pareodinia spp. 62 rigaudella filamentosa 63 rhyncodiniopsis spp. 64 barbatacysta verrucosa 65 endoscrinium galeritum 66 scriniodinium crystallinium 67 leptodinium subtile 68 escharisphaeridia pocockii 69 scriniodinium spp. 70 pareodinia scopaeus 71 apteodinium cf. nuciforme 72 scriniodinium inritibilum 73 tubotuberella cf. dangeardii 74 endoscrinium cf. galeritum 75 meiourogonyaulax spp. 76 epiplosphaera spp. 77 dingodinium jurassicum 78 tubotuberella cf. apatela 79 escarisphaeridia erythrocoma 80 sentusidinium sp. e (fensome 1979) 81 valensiella cf. dictydia 82 tenua cf. hystrix 83 sirmiodiniopsis spp. 84 escharispahaeria spp. 85 pareodinia borealis 86 rhynchodiniopsis sp. (cf. "machaera") 87 atopodinium cf. haromense 88 rhynchodiniopsis spp. 89 avellodinium spp. 90 systematophora spp. 91 sentusidinium myriatrichum 92 prolixosphaeridium granulosum 93 cribroperidinium granuligerum 94 endoscrinium luridum 95 circulodinium distinctum 96 scriniodinium irregulare 97 glossodinium dimorphum 98 adnatosphaeridium spp. > 50 specim ens 20–50 specim ens 5–19 specim ens 2–4 specim ens 1 specim en g e u s b ulletin no 5.pm d 29-10-2004, 11:14 83 84 jason chronozone and trichodinium scarburghense appears in the p. athleta chronozone (milner & piasecki 1996; piasecki 1996). age. the age of the upper sandstone unit of the spath plateau member, pelion formation, is therefore early to late callovian, equivalent to the p. koenigi – p. athleta chronozones (fig. 4). depositional environment. the organic matter is dominated by terrestrial palynomorphs and debris. the proportion of brown and black lath-shaped woody material increases upwards until it completely dominates the organic content in the upper pelion formation. the upwards increase and dominance of woody material suggests deposition in the lower shoreface environment in front of a prograding shoreline. payer dal formation the payer dal formation was analysed from two localities at steensby bjerg, along the sortelv (locality 2) and gulelv (locality 1) rivers (figs 2, 4, 6, 7). the formation is characterised by frequent rigaudella aemula, rhynchodiniopsis cladophora, sirmiodinium grossii and gonyaulacysta jurassica. new, stratigraphically characteristic species appear in the lower part of the formation at sortelv: wanaea digitata, rigaudella aemula and leptodinium subtile. higher in the formation at both localities further stratigraphically significant species appear: scriniodinium crystallinum, endoscrinium galeritum, chytroeisphaeridia chytroeides, rhynchodiniopsis sp., prolixosphaeridium granulosum and dingodinium jurassicum. correlation. the dinoflagellate assemblage represents a characteristic lower to middle oxfordian assemblage with frequent rigaudella aemula, scrinidinium crystallinum and endoscrinium galeritum, as known from the jurassic succession elsewhere in east greenland such as in milne land (piasecki 1996). this lower– middle oxfordian assemblage in milne land reaches close to the top of the middle oxfordian before gradual replacement by an upper oxfordian assemblage. wanaea spp. occurs only to the top of the lower alphabetical species list 98 adnatosphaeridium spp. 19 aldorfia aldorfensis 54 ambonosphaera calloviense 71 apteodinium cf. nuciforme 87 atopodinium cf. haromense 53 atopodinium haromense 21 atopodinium polygonale 9 atopodinium spp. 89 avellodinium spp. 23 barbatacysta creberbarbata 64 barbatacysta verrucosa 52 chlamydophorella ectotabulata 3 chytroeisphaeridia cerastes 5 chytroeisphaeridia hyalina 95 circulodinium distinctum 93 cribroperidinium granuligerum 17 ctenidodinium thulium 33 cyclopsiella spp. 77 dingodinium jurassicum 47 durotrigia spp. 74 endoscrinium cf. galeritum 65 endoscrinium galeritum 94 endoscrinium luridum 76 epiplosphaera spp. 79 escarisphaeridia erythrocoma 8 escharispahaeria laevigata 84 escharispahaeria spp. 6 escharisphaeridia cf. pocockii 68 escharisphaeridia pocockii 45 escharisphaeridia rudis 43 fromea tornalis 97 glossodinium dimorphum 60 gonyaulacysta eisenackii 15 gonyaulacysta cf. helicoidea 25 gonyaulacysta jurassica 56 gonyaulacysta spp. 22 kallosphaeridium hypornatum 32 kallosphaeridium praussii 46 kallosphaeridium spp. 1 limbicysta bjaerkei 41 lithodinia cf. spongiosa 44 lithodinia jurassica 20 lithodinia planoseptata 30 lithodinia spongiosa 18 lithodinia spp. 49 lithodinia valensi 75 meiourogonyaulax spp. 50 mendicodinium "granulatum" 37 mendicodinium groenlandicum 48 nannoceratopsis plegas var. dictyornata 13 nannoceratopsis pellucida 2 paraevansia brachythelis 28 paragonyaulacysta retiphragmata 12 pareodinia "granulata" 85 pareodinia borealis 38 pareodinia ceratophora 31 pareodinia halosa 51 pareodinia pachyceras 16 pareodinia prolongata 70 pareodinia scopaeus 61 pareodinia spp. 57 pareodinia stegasta 92 prolixosphaeridium granulosum 4 rhynchodiniopsis cf. cladophora 86 rhynchodiniopsis sp. (cf. "machaera") 26 rhynchodiniopsis cladophora 88 rhynchodiniopsis spp. 63 rhynchodiniopsis spp. 62 rigaudella filamentosa 66 scriniodinium crystallinium 72 scriniodinium inritibilum 96 scriniodinium irregulare 69 scriniodinium spp. 10 sentusidinium cf. pelionense 91 sentusidinium myriatrichum 55 sentusidinium pelionense 34 sentusidinium sp. d (fensome 1979) 80 sentusidinium sp. e (fensome 1979) 36 sentusidinium spp. 83 sirmiodiniopsis spp. 11 sirmiodinium grossii 35 surculosphaeridium spp. 90 systematophora spp. 82 tenua cf. hystrix 59 trichodinium scarburghense 78 tubotuberella cf. apatela 73 tubotuberella cf. dangeardii 40 tubotuberella dangeardii 58 tubotuberella eisenackii 27 tubotuberella spp. 81 valensiella cf. dictydia 7 valensiella dictydia 14 valensiella ovula 42 valensiella sp. (fensome 1979) 29 valensiella spp. 39 veryhachium spp. 24 algae indet. 67 leptodinium subtile previous page: fig. 6. distribution chart of dinoflagellate cysts in the jurassic succession at locality 1, steensby bjerg, northern hold with hope. the first appearances of species are stratigraphically arranged. alphabetical species list given above. geus bulletin no 5.pmd 29-10-2004, 11:1484 85 hold with hope, locality 2, sortelv 20 sa m pl es h ei gh t 34.00 32.00 28.50 25.00 9.00 g g u s am pl e no . 433859 433159 443158 433157 433858 la te ju ra ss ic o xf or di an pa ye r d al f or m at io n 1 s cr in io di ni um cf . i nr iti bi lu m 2 c hy tr oe is ph ae rid ia c er as te s 3 w an ae a sp p. 4 r ig au de lla fi la m en to sa 5 s te ph an el yt ro n sp p. 6 a to po di ni um h ar om en se 7 g on ya ul ac ys ta ju ra ss ic a 8 s irm io di ni um g ro ss ii 9 r hy nc ho di ni op si s cl ad op ho ra 10 tr ic ho di ni um s ca rb ur gh en se 11 e nd os cr in iu m g al er itu m 12 le pt od in iu m s ub til e 13 r ig au de lla a em ul a 14 a to po di ni um sp p. 15 p ar eo di ni a "g ra nu la ta " 16 s ur cu lo sp ha er id iu m sp p. 17 c irc ul od in iu m d is tin ct um 18 tu bo tu be re lla a pa te la 19 w an ae a di gi ta ta 20 p ar eo di ni a bo re al is 21 e sc ha ris ph ae rid ia la ev ig at a 22 v al en si el la d ic ty di a 23 s cr in io di ni um c ry st al lin iu m 24 a m bo no sp ha er a ca llo vi an a 25 p ar eo di ni a ce ra to ph or a 26 c hy tr oe is ph ae rid ia c hy tr oe oi de s 27 p ro lix os ph ae rid iu m g ra nu lo su m 28 e sc ha ris ph ae rid ia sp p. alphabetical species list 24 ambonosphaera calloviana 6 atopodinium haromense 14 atopodinium spp. 2 chytroeisphaeridia cerastes 26 chytroeisphaeridia chytroeoides 17 circulodinium distinctum 11 endoscrinium galeritum 21 escharisphaeridia laevigata 28 escharisphaeridia spp. 7 gonyaulacysta jurassica 12 leptodinium subtile 15 pareodinia "granulata" 20 pareodinia borealis 25 pareodinia ceratophora 27 prolixosphaeridium granulosum 9 rhynchodiniopsis cladophora 13 rigaudella aemula 4 rigaudella filamentosa 1 scriniodinium cf. inritibilum 23 scriniodinium crystallinium 8 sirmiodinium grossii 5 stephanelytron spp. 16 surculosphaeridium spp. 10 trichodinium scarburghense 18 tubotuberella apatela 22 valensiella dictydia 19 wanaea digitata 3 wanaea spp. m et re s a ge st ag e fo rm at io n 30 10 20–50 specimens 5–19 specimens 2–4 specimens 1 specimen fig. 7. distribution chart of dinoflagellate cysts in the jurassic succession at locality 2, sortelv, western steensby bjerg, hold with hope. the first appearance of species is stratigraphically arranged. g e u s b ul le tin n o 5. pm d 29 -1 020 04 , 1 1: 14 85 86 oxfordian in milne land, whereas leptodinium subtile rarely occurs below the middle oxfordian and prolixosphaeridium granulosum does not occur below the upper oxfordian. the payer dal formation at locality 1 is therefore considered middle to late oxfordian in age, and the presence of wanaea digitata, wanaea sp. and trichodinium scharburghense in the assemblage is due to reworking. in milne land, the appearance of leptodinium subtile in the c. tenuiserratum chronozone coincides with the gradual disappearance of rigaudella spp., and the following appearances of dingodinium jurassicum and prolixosphaeridium granulosum in the a. glosense chronozone. a corresponding sequence of events in the steensby bjerg succession indicates a middle–upper oxfordian succession, c. tenuiserratum – a. glosense chronozones. consequently, a significant late callovian – earliest middle oxfordian hiatus occurs between the pelion and payer dal formations. however, several samples in the boundary interval (c. 30 m thick) were barren of dinoflagellate cysts and parts of the succession were therefore not dated. the dinoflagellate cysts, which are considered reworked, indicate that lower oxfordian sediments have been present in the region. age. the age of the payer dal formation is middle– late oxfordian, equivalent to the c. tenuiserratum – a. glosense chronozones based on dinoflagellate cysts. ammonites in the overlying bernbjerg formation support this age of the uppermost payer dal formation as they indicate the a. glosense chronozone. depositional environment. the organic matter is dominated by terrestrial palynomorphs and debris, and the proportion of brown and black lath-shaped woody material is high in the payer dal formation. the organic content suggests deposition in a lower shoreface environment. bernbjerg formation the bernbjerg formation is represented by a few samples from the lower and upper parts of the formation at steensby bjerg (locality 1). the bernbjerg formation contains abundant sirmiodinium grossii and gonyaulacysta jurassica. in the lower levels of the formation, the presence of abundant leptodinium subtile is combined with the appearance of paragonyaulacysta borealis, rhynchodiniopsis sp. and tenua cf. hystrix. the assemblage is very similar to the assemblage in the upper payer dal formation partly due to the continued presence of endoscrinium galeritum and scriniodinium crystallinum. the stratigraphically important taeniophora sp. / adnatosphaeridium sp. (informal name ‘a. hartzii’ in: piasecki 1980) appears in the uppermost sample from the formation. correlation. the continued presence of endoscrinium galeritum and scriniodinium crystallinum from the payer dal formation below, and the absence of taeniophora sp. / adnathosphaeridium sp. (‘a. hartzii’) correlates with the lower upper oxfordian, a. glosense chronozone, by comparison to dinoflagellate floras from milne land (piasecki 1996). this is in accordance with abundant ammonites of the amoeboceras glosense zone in these strata, and with the absence of the uppermost oxfordian dinoflagellate cyst species that appear above. the upper part of the bernbjerg formation contains abundant gonyaulacysta jurassica and sirmiodinium grossii in combination with adnatosphaeridium sp. (‘a. hartzii’), cribroperidinium granuligerum, sciniodinium irregulare, glossodinium cf. dimorphum, endoscrinium luridum and prolixosphaeridium granulosum. the composite dinoflagellate cyst flora indicates an upper oxfordian to lowermost kimmeridgian succession, a. serratum – p. baylei chronozones. the presence of avellodinium spp. in the uppermost sample could indicate the lowermost kimmeridgian a. mutabilis chronozone, but this is not supported by any other stratigraphical diagnostic species such as perisseiasphaeridium pannosum (piasecki 1996; piasecki & stemmerik 2004, this volume). age. the age of the bernbjerg formation is late oxfordian – earliest kimmeridgian, equivalent to the a. glosense – p. baylei chronozones based on dinoflagellate cysts. ammonites in the lower part of the bernbjerg formation indicate the a. glosense chronozone and confirm the late oxfordian age for this part of the formation. depositional environment. the organic content is dominated by terrestrial sporomophs and woody material but dinoflagellate cysts occur frequently. a depositional environment of lower shoreface to open shelf is interpreted on this basis. geus bulletin no 5.pmd 29-10-2004, 11:1486 87 correlations the pelion formation on northern hold with hope comprises two main units, a lower sandstone unit followed by mudstones and heterolithic sandstones of the spath plateau member. the same overall pattern occurs in the pelion formation on store koldewey at ravn pynt (piasecki et al. 2004, this volume). however, on store koldewey, the lower sandstone unit is older (bathonian) than the lower sandstone unit on hold with hope. the mudstone and overlying sandstone on store koldewey are basically of the same early callovian age as the lowermost spath plateau member on hold with hope (c. apertum – p. koenigi chronozones). the marine flooding represented by deposition of this mudstone can be traced from milne land and jameson land in the south (p7 – third order sequence; engkilde & surlyk 2003) to hold with hope and store koldewey in the north. the payer dal formation is defined on kuhn ø where it comprises two units that are of early–middle oxfordian age and early late oxfordian age (alsgaard et al. 2003). on hold with hope, the exposure of the oldest part of the payer dal formation at sortelv (fig. 2; locality 2) is limited by a fault, and older strata may be present in the subsurface. however, no strata of early oxfordian age have been recorded here, and the age of the payer dal formation on hold with hope is middle–late oxfordian, partly corresponding to the upper part of this formation on kuhn ø. sedimentation of fine-grained sand and mudstone of the bernbjerg formation began in the a. glosense chron on hold with hope as in wollaston forland to the north (surlyk 1977). conclusions the combined biostratigraphical dataset from ammonites and dinoflagellate cysts dates the stratigraphical range of the lithological units with a high degree of precision (figs 3, 4). however, the extent of non-depositional or erosional hiati in or between the units cannot be determined with the same certainty due to the limited number of productive samples. the ‘basal sandstone unit’ of the pelion formation ranges from the uppermost c. apertum chronozone to the lower p. koenigi chronozone (figs 3, 4). the age is therefore early callovian. the dinoflagellate assemblages show no indication of a break in sedimentation so the succession is considered complete. the spath plateau member of the pelion formation comprises the p. koenigi, k. jason and p. athleta chronozones (fig. 3). the age is therefore early to late callovian, but a part of the succession occurs above the highest productive sample and may therefore be younger. a considerable hiatus is present between the pelion formation and the overlying payer dal formation. however, the exact stratigraphical position of the unconformity and the extent of the hiatus cannot be determined precisely, because no productive samples were recovered from the boundary interval. the available data suggest a hiatus that comprises most of the late callovian, early oxfordian and earliest middle oxfordian. the payer dal formation ranges from the c. tenuiserratum to the a. glosense chronozones, and the age is consequently middle to late oxfordian (fig. 3). the dinoflagellate assemblages indicate no break in deposition at the boundary to the bernbjerg formation, and the a. glosense chronozone is also identified in the basal bernbjerg formation on the basis of ammonites. the jurassic succession and the bernbjerg formation are limited upwards by pre-barremian, cretaceous erosion, and the highest samples are referred to the a. rosenkrantzi – p. baylei chronozones at the oxfordian–kimmeridgian boundary (fig. 3). the age of the bernbjerg formation is thus late oxfordian, possibly earliest kimmeridgian. the jurassic succession on northern hold with hope correlates well with the corresponding jurassic successions towards the north and the south, but appears more fragmented compared to these successions. a boreal bathonian (bajocian–bathonian) succession has been deposited in this region, at least partly, but was removed by later erosion as indicated by the reworked ammonite of the p. pompeckji zone. the previous presence of a lower oxfordian succession is similarly indicated by reworked dinoflagellate cysts. the magnitude of the hiatus below the payer dal formation is late callovian – middle oxfordian. acknowledgements the present biostratigraphic study was supported by the carlsberg foundation (carlsbergfondet ans. 980089/ 20-262). john h. callomon and peter alsen are thanked for identification of ammonites from hold with hope. the authors are grateful to jan jansonius and susanne feist-burkhardt for useful and constructive review comments. geus bulletin no 5.pmd 29-10-2004, 11:1487 88 references alsgaard, p.c., felt, v.l., vosgerau, h. & surlyk, f. 2003: the jurassic of kuhn ø, north-east greenland. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 865–892. bailey, d. & hogg, n.m. 1995: fentonia bjaerkei gen. et comb. nov.; transfer from parvocysta bjaerke 1980. journal of micropalaeontology 14(1), 58 pp. callomon, j.h. 1993: the ammonite succession in the middle jurassic of east greenland. bulletin of the geological society of denmark 40, 83–113. donovan, d.t. 1957: the jurassic and cretaceous systems in east greenland. meddelelser om grønland 155(4), 1–214. engkilde, m. & surlyk, f. 2003: shallow marine syn-rift sedimentation: middle jurassic pelion formation, jameson land, east greenland. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 813–863. fensome, r.a. 1979: dinoflagellate cysts and acritarchs from the middle and upper jurassic of jameson land, east greenland.bulletingrønlands geologiske undersøgelse 132, 98 pp. kelly, s.r.a., whitham, a.g., koraini, a.m. & price, s.p. 1998: lithostratigraphy of the cretaceous (barremian–santonian) hold with hope group, ne greenland. journal of the geological society (london) 155(6), 993–1008. larsen, m., piasecki, s., preuss, t., seidler, l., stemmerik, l., therkelsen, j. & vosgerau, h. 1997: petroleum geological activities in east greenland in 1997. geology of greenland survey bulletin 180, 35–42. milner, p.s. & piasecki, s. 1996: boreal middle jurassic dinoflagellate cyst stratigraphy of jameson land, east greenland. in: piasecki, s. et al. (eds): formation of source and reservoir rocks in a sequence stratigraphic framework, jameson land, east greenland. energy research programme efp-93, projects 1313/93-0010 and 0017. danmarks og grønlands geologiske undersøgelse rapport 1996/30, vol. i & ii, 46 pp. piasecki, s. 1980: middle to late jurassic dinoflagellate cyst stratigraphy from milne land and jameson land (east greenland) correlated with ammonite stratigraphy, 167 pp. unpublished ph.d. thesis, university of copenhagen, denmark. piasecki, s. 1996: boreal dinoflagellate cyst stratigraphy of middle to upper jurassic sediments of milne land, east greenland. in: piasecki, s. et al. (eds): formation of source and reservoir rocks in a sequence stratigraphic framework, jameson land, east greenland. energy research programme efp93, projects 1313/93-0010 and 0017. danmarks og grønlands geologiske undersøgelse rapport 1996/30, vol. i & ii, 100 pp. piasecki, s. & stemmerik, l. 2004: jurassic dinoflagellate cysts from hochstetter forland, north-east greenland. in: stemmerik, l. & stouge, s. (eds): the jurassic of north-east greenland. geological survey of denmark and greenland bulletin 5, 89–97 (this volume). piasecki, s., callomon, j.h. & stemmerik, l. 2004: jurassic dinoflagellate cyst stratigraphy of store koldewey, north-east greenland. in: stemmerik, l. & stouge, s. (eds): the jurassic of north-east greenland. geological survey of denmark and greenland bulletin 5, 99–112 (this volume). poulsen, n.e., gudmundsson, l., hansen, j.m. & husfelt, y. 1990: palynological preparation techniques, a new macerationtank-method and other modifications. danmarks geologiske undersøgelse serie c 10, 22 pp. price, s.p. & whitham, a.g. 1997: exhumed hydrocarbon traps in east greenland: analogs for the lower–middle jurassic play of northwest europe. american association of petroleum geologists bulletin 81, 196–221. stemmerik, l., clausen, o.r., korstgård, j., larsen, m., piasecki, s., seidler, l., surlyk, f. & therkelsen, j. 1997: petroleum geological investigations in east greenland: project ‘resources of the sedimentary basins of north and east greenland’. geology of greenland survey bulletin 176, 29–38. surlyk, f. 1977: stratigraphy, tectonics and palaeogeography of the jurassic sediments of the areas north of kong oscars fjord, east greenland. bulletin grønlands geologiske undersøgelse 123, 56 pp. surlyk, f. 1978: mesozoic geology and palaeogeography of hochstetter forland, east greenland. bulletin of the geological society of denmark 27, 73–87. surlyk, f., clemmensen, l.b. & larsen, h.c. 1981: post-palaeozoic evolution of the east greenland continental margin. in: kerr, j.w. & fergusson, a.j. (eds): geology of the north atlantic borderlands. canadian society of petroleum geologists memoir 7, 611–645. vischer, a. 1943: die postdevonische tektonik von ostgrönland zwischen 74º und 75ºn. br. kuhn ø, wollaston forland, clavering ø und angrenzende gebiete. meddelelser om grønland 133(1), 195 pp. vosgerau, h., larsen, m., piasecki, s. & therkelsen, j. 2004: a new middle–upper jurassic succession on hold with hope, north-east greenland. in: stemmerik, l. & stouge, s. (eds): the jurassic of north-east greenland. geological survey of denmark and greenland bulletin 5, 51–71 (this volume). geus bulletin no 5.pmd 29-10-2004, 11:1488 geological survey of denmark and greenland bulletin 42, 2018, 127-131 127 igneous intrusions in the cored upper jurassic succession of the blokelv-1 borehole, jameson land basin, east greenland lotte melchior larsen the fully cored upper jurassic succession in the blokelv-1 borehole in the jameson land basin, east greenland, is intersected by igneous intrusions at four levels; the intrusions comprise a c. 15 cm thick dyke and three sills with thicknesses of 0.7, 1.2 and 1.9 m. the sills consist of fine-grained, sparsely plagioclase-olivine-phyric basalt with chilled contacts to the sediments. analyses of two sills gave very similar results. the sills are tholeiitic basalts with compositions similar to the main group of dykes and sills in the jameson land basin, and the blokelv-1 sills are thus considered to belong to this group which has been dated at c. 53 ma. the intrusions form part of a 55–51 ma suite of tholeiitic basalt intrusions that was emplaced over an area extending for over 500 km north-to-south within the sedimentary basins of east and north-east greenland. keywords: east greenland, dykes, cenozoic __________________________________________________________________________ l.m.l., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: lml@geus.dk the fully cored blokelv-1 borehole was drilled in 2008 through upper jurassic sediments in the jameson land basin, east greenland (bojesen-koefoed et al. 2009), in order to study the sedimentary succession (fig. 1a). the sediments in jameson land are intruded by many cenozoic dykes and sills (e.g. noe-nygaard 1976; hald & tegner 2000), and one of the site selection criteria was to minimise the risk of encountering thick igneous intrusions during drilling. although major intrusions were avoided, the 233.8 m succession in the blokelv-1 core is cut by intrusive igneous rocks at four levels with a combined thickness of 4.1 m (fig. 1b). the purpose of this paper is to present descriptions and analyses of these intrusions and compare them with other cenozoic dyke and sill intrusions in northern east greenland. intrusions in the blokelv-1 core the core is cut by igneous intrusions at four levels in the upper half of the section: 102.04–100.1 m (thickness 1.9 m), 56.4–55.2 m (thickness 1.2 m), 27.10–26.40 m (thickness 0.70 m) and 7.35–7.05 m (thickness 0.3 m). the uppermost intrusion has oblique boundary contacts, dipping at 60°, and is accordingly described as a dyke; its true thickness must be c. 15 cm. the three lower intrusions show boundary contacts that are broadly parallel to bedding in the host rock and thus appear to be sills. the two thickest sills have caused prominent alteration of the surrounding sediments (see olivarius et al. 2018, this volume). the thin dyke uppermost in the section (fig. 1b) is thoroughly altered and was not studied further. the remaining three sills are lithologically similar and consist of fine-grained, sparsely plagioclase-olivine-phyric basalt. at the chilled contacts, they are very fine-grained to aphanitic, altered, and cut by carbonate veins. in the middle sill, a fracture is filled with biodegraded oil (bojesenkoefoed et al. 2018, this volume). © geus, 2018. geological survey of denmark and greenland bulletin 42, 127–131. available at: www.geus.dk/bulletin42 mailto:lml@geus.dk http://www.geus.dk/bulletin42 128128 petrography the 1.9 m thick lower sill (fig. 1b, 102.04–100.1 m) has a lower chilled margin that is aphanitic with many plagioclase microlites, sparse 0.1–0.2 mm plagioclase microphenocrysts, and a few plagioclase-olivine glomerocrysts with up to 2 mm plagioclase laths and 0.8 mm olivine crystals. plagioclase is fresh but the olivine is completely altered. the groundmass is extensively replaced by carbonate, but plagioclase phenocrysts are fresh. the rock is cut by 0.1–0.5 mm wide veins of ankerite with patches of pyrite which cut both sill and sandstone at the contact (fig. 2; see olivarius et al. 2018, this volume). the veins are thickest at the contact. the veins in the sandstone appear to fill tension cracks. the 1.2 m thick middle sill (fig. 1b, 56.4–55.2 m) has a very fine-grained upper chilled margin with sparse 0.5 mm plagioclase microphenocrysts and tiny <0.5 mm euhedral olivine crystals; all olivine crystals are altered to clay. the groundmass is intersertal with numerous plagioclase microlites. the central part of the sill is finegrained with sparse <1 mm plagioclase phenocrysts and sparse c. 0.5 mm fresh olivine microphenocrysts, often assembled in glomerocrysts. the groundmass is intergranular with plagioclase, clinopyroxene, fe-ti oxide, olivine, and abundant mesostasis. the 0.7 m thick upper sill (fig. 1b, 27.10–26.40 m) is a fine-grained, nearly aphyric rock with scattered vugs filled with colourless minerals. the lower contact is not preserved. the rock becomes slightly finer grained towards the upper contact and there is possibly a thin glass chill at the top. there is no visible influence on the overlying sediments. 70°30'n 50 km blokelv-1 mudstone sandstone a igneous intrusion 25°w 24°w 23°w 22°w 71°30'n 71°n 24°w 23°w 22°w 10 depth (m) 20 30 40 50 60 70 80 90 100 110 b h urry in le t jameson land liverpool land bk s lml 1 s c o r e s b y s u n d upper jurassic – lower cretaceous middle jurassic upper triassic – lower jurassic devonian – middle triassic crystalline rock normal fault fig. 1. a: map of central and southern jameson land showing the location of the blokelv-1 borehole (70°45.305́ n, 23°40.430´w, wgs84 coordinates); inset shows the location of the study area in east greenland. bk: blosseville kyst. s: shannon. b: simplified log of the upper part (0–110 m) of the blokelv-1 core (ggu no. 511101). 129 chemical compositions two samples from the centres of the lower and middle sills have been analysed for major and trace elements. major elements were analysed by n. odling at university of edinburgh by x-ray fluorescence spectrometry (xrf) and procedures as described by fitton et al. (1998). trace elements were analysed in geus’ rock geochemical laboratory using a perkinelmer elan 6100 drc quadrupole inductively coupled plasma mass spectrometer (icp-ms). sample dissolution followed a modified version of the procedure used by turner et al. (1999) and ottley et al. (2003). calibration was done using two certified ree solutions and three international reference standards. results for reference samples processed and run simultaneously with the unknowns are normally within 5% of the reference value for most elements with concentrations >0.1 ppm (results are shown in table 1). the two sills consist of tholeiitic basalt with 6–7 wt% mgo and 2.1–2.3 wt% tio2. losses on ignition are low and the samples appear to be fresh. measured values of tantalum (ta) are high (1.1–1.4 ppm) where only c. 0.8 ppm ta is expected; this may be contamination from the tungsten carbide crushing vessel and the data are not included in table 1. the lower sill has 5.6 ppm pb which is very high, indicating contamination with pb either during emplacement or from the drilling process; there is no evidence of additional contamination. 5 mm fig. 2. lower contact towards sandstone of the 1.9 m thick lower sill at 102.04 m in the core. two ankerite veins cross both sill and sandstone. note plagioclase-olivine glomerocryst in the very fine-grained basalt matrix between the two veins. thin section 511101.246; plane-polarised light. major elements, wt% (xrf analyses) sio2 47.84 48.43 48.34 48.30 tio2 2.30 2.11 2.35 2.33 al2o3 13.71 12.87 13.79 13.80 fe2o3 14.07 12.95 14.07 14.07 mno 0.21 0.26 0.20 0.20 mgo 6.85 6.24 6.99 6.96 cao 11.23 12.08 11.52 11.47 na2o 2.23 2.03 2.39 2.37 k2o 0.22 0.40 0.29 0.30 p2o5 0.20 0.19 0.24 0.23 loi 0.47 1.84 -0.27 0.14 sum 99.33 99.40 99.91 100.17 trace elements, ppm (icp-ms analyses) sc 38 36 40 38 v 374 345 371 362 cr 190 179 259 249 co 55 50 53 52 ni 99 90 121 109 cu 240 218 256 249 zn 108 100 114 115 ga 21.0 20.0 rb 3.4 8.7 5.3 6.3 sr 217 230 237 228 y 31.1 30.5 32.1 31.7 zr 145 135 159 159 nb 12.1 11.7 12.6 12.8 cs 0.52 0.20 0.48 0.12 ba 130 110 71.6 71.9 la 10.5 11.6 10.7 10.9 ce 27.2 29.9 28.0 28.0 pr 4.10 4.27 4.16 4.17 nd 19.7 20.1 19.1 19.2 sm 5.09 5.01 5.15 5.27 eu 1.68 1.64 1.77 1.74 gd 6.07 5.79 5.42 5.30 tb 0.99 0.91 0.94 0.93 dy 5.78 5.45 5.49 5.36 ho 1.14 1.07 1.13 1.10 er 3.09 2.95 2.91 2.87 tm 0.47 0.44 0.43 0.43 yb 2.77 2.69 2.65 2.61 lu 0.41 0.39 0.36 0.38 hf 3.77 3.47 3.60 3.57 ta 0.84 0.84 pb 0.91 5.61 1.29 1.71 th 0.84 1.25 0.9 0.89 u 0.26 0.29 0.31 0.31 depth (m) 55.2–56.4 100.1–102 ene dyke thin sill ggu no 511101.230 511101.229 407203 403021 table 1. chemical analyses of two sills in the blokelv core, with comparisons from jameson land total iron is reported as fe2o3. loi is loss on ignition. data for jameson land from hald & tegner (2000). middle sill lower sill jameson land lml table 1 130130 discussion the two analysed sills have very similar compositions and are considered to have been intruded during the same magmatic event. the tholeiitic basalt represents a magma type that is known from widespread sills and dykes in the jameson land basin (larsen et al. 1989; hald & tegner 2000). hald & tegner (2000) recognised five different magma types represented by sills and dykes, and by far the most common group is the one found in the blokelv sills. this group was called the ‘high-ti group’ by hald & tegner (2000), but as the ti contents are not high, it is referred to here as the ‘main group’. figure 3 shows geochemical patterns for the blokelv sills compared with similar patterns for the jameson land sills and dykes. the close similarity of the blokelv sills with the main group of tholeiitic sills and dykes in jameson land is clear. the blokelv intrusions are poorly suited for 39ar/40ar dating because of the low k2o content and few and small plagioclase phenocrysts. however, hald & tegner (2000) dated a sill and a dyke from the main group by the 39ar/40ar method. the sill yielded a 5-point isochron age of 52.7 ± 1.2 ma, and the dyke yielded a four-point isochron age of 53.3 ± 1.4 ma; the two ages are within the uncertainty of each other (the ages are here recalculated to an age of 28.201 ma for the fish canyon tuff standard). it is therefore most probable that the two blokelv sills were emplaced at c. 53 ma. the intrusions in the jameson land basin were emplaced after the plateau lavas of the blosseville kyst at 56.4–55.3 ma (storey et al. 2007) and after or just concomitantly with the plateau lavas in north-eastern greenland at 56–53 ma (larsen et al. 2014). they are within the age range of 55–51 ma obtained for tholeiitic sills and dykes intruded into the sediments from jameson land in the south to the island of shannon in the north (hald & tegner 2000; larsen et al. 2014). most intrusion ages are in the interval 54–52 ma and magma production at that time must have been very extensive. these intrusions cover a stretch of least 500 km which is close to the entire onshore extent of the mesozoic basins. the intrusions are older than the igtertivâ formation at kap dalton on the blosseville kyst which comprises two parts dated at 49.1 ± 0.5 ma and 43.8 ± 1.1 ma (larsen et al. 2013). larsen et al. (2013) found significant geochemical differences between the igtertivâ formation basalts and the underlying 55 ma lavas of the skrænterne formation, in particular in the rare-earth element (ree) ratios. as seen in fig. 4, the intrusions at 55–51 ma retained the geochemical characteristics of the older plateau lavas, indicating that the conditions of magma generation were unchanged, probably mainly governed by the relatively thick lithosphere beneath the continent away from the developing oceanic rift (hald & tegner 2000). conclusions the two analysed sills in the blokelv-1 core are compositionally similar to the main group (high-ti group) of tholeiitic basalt sills and dykes that occur frequently in the jameson land basin, of which two have been dated at c. 53 ma. the blokelv sills are considered to belong to this group. the group shares trace element characteristics 1 10 100 1000 1 10 100 1000 la ce pr nd pm sm eu gd tb dy ho er tm yb lu rb ba th u nb ta k la ce pb pr sr nd sm zr hf eu ti tb dy y ybgdp sa m pl e/ c ho nd rit e sa m pl e/ pr im iti ve m an tle alkaline main group blokelv sills low-ti lml 2lml3 fig. 3. multi-element patterns for the blokelv sills compared with sills and dykes in the jameson land basin (data from hald & tegner 2000). alkaline and low-ti basalt are two other basalt groups defined by hald & tegner (2000). note the close similarity between the main group of intrusions and the blokelv sills. the high pb in one of the blokelv samples must be due to contamination. 131 with the older blosseville kyst lavas and the 55−51 ma tholeiitic sills and dykes in east greenland, but not with the younger 49−54 ma igtertivâ formation lavas. the older magmas were probably generated under similar conditions beneath relatively thick lithosphere. acknowledgements the referees, godfrey fitton and christian tegner, are thanked for their constructive comments. references bojesen-koefoed, j.a., bjerager, m. & piasecki, s. 2009: shallow core drilling and petroleum geology related field work in east and north-east greenland 2008. geological survey of denmark and greenland bulletin 17, 53–56. bojesen-kofoed, j., bjerager, m., nytoft, h.p., petersen, h.i., piasecki, s. & pilgaard, a. 2018: petroleum potential of the upper jurassic hareelv formation, jameson land, east greenland. in: ineson, j. & bojesen-koefoed, j.a. (eds): petroleum geology of the upper jurassic – lower cretaceous of east and north-east greenland: blokelv-1 borehole, jameson land basin. geological survey of denmark and greenland bulletin 42, 85–113 (this volume). fitton, j.g., saunders, a.d., larsen, l.m., hardarson, b.s. & norry, m.j. 1998: volcanic rocks from the southeast greenland margin at 63°n: composition, petrogenesis and mantle sources. in: saunders, a.d., larsen, h.c. & wise, s.w., jr. (eds): proceedings of the ocean drilling program, scientific results 152, 331−350. college station, tx. hald, n. & tegner, c. 2000: composition and age of tertiary sills and dykes, jameson land basin, east greenland: relation to regional flood volcanism. lithos 54, 207–233. larsen, l.m., watt, w.s. & watt. m. 1989: geology and petrology of the lower tertiary plateau basalts of the scoresby sund region, east greenland. bulletin grønlands geologiske undersøgelse 157, 164 pp. larsen, l.m., pedersen, a.k., sørensen, e.v., watt, w.s. & duncan, r.a. 2013: stratigraphy and age of the eocene igtertivâ formation basalts, alkaline pebbles and sediments of the kap dalton group in the graben at kap dalton, east greenland. bulletin of the geological society of denmark 61, 1–18. larsen, l.m., pedersen, a.k., tegner, c. & duncan, r.a. 2014: eocene to miocene igneous activity in ne greenland: northward younging of magmatism along the east greenland margin. journal of the geological society (london) 171, 539–553. noe-nygaard, a. 1976: tertiary igneous rocks between shannon and scoresby sund, east greenland. in: escher, a. & watt, w.s. (eds): geology of greenland, 386–402. copenhagen: geological survey of greenland. olivarius, m., weibel, r., schovsbo, n.h., olsen, d. & kjøller, c. 2018: diagenesis of upper jurassic sandstones of the blokelv-1 core in the jameson land basin, east greenland. in: ineson, j. & bojesen-koefoed, j.a. (eds): petroleum geology of the upper jurassic – lower cretaceous of east and north-east greenland: blokelv-1 borehole, jameson land basin. geology of denmark and greenland survey bulletin 42, 65 –84 (this volume). ottley, c.j., pearson, d.g. & irvine, g.j. 2003: a routine method for the dissolution of geological samples for the analysis of ree and trace elements via icp-ms. in: holland, j.g. & tanner, s.d. (eds): plasma source mass spectrometry: applications and emerging technologies, 221–230. cambridge: royal society of chemistry. storey, m., duncan, r.a. & tegner, c. 2007: timing and duration of volcanism in the north atlantic igneous province: implications for geodynamics and links to the iceland hotspot. chemical geology 241, 264–281. turner, s.p., platt, j.p., george, r.m.m., kelly, s.p., pearson, d.g. & nowell, g.m. 1999: magmatism associated with orogenic collapse of the betic–alboran domain, se spain. journal of petrology 40, 1011–1036. 1.0 1.5 2.0 2.5 0.0 0.5 1.0 1.5 2.0 2.5 (la/sm)n (g d/ lu ) n igtertivâ fm lavas skrænterne fm lavas ne greenland sills and dykes jameson land sills and dykes blokelv intrusions, jameson land 55–51 ma 49–44 ma lml 4 fig. 4. rare-earth element (ree) ratios for older (55–51 ma) lavas and intrusions (blosseville kyst to shannon) and younger (49–44 ma) igtertivâ formation lavas on blosseville kyst, east greenland. data from hald & tegner (2000), larsen et al. (2013), larsen et al. (2014) and unpublished geus data (2008–2010). a few crustally contaminated samples are not plotted. la/sm is the light ree ratio and gd/lu is the heavy ree ratio; n designates chondrite-normalised concentrations. _________________________________________________________________________________________ manuscript received 17 december 2015; revision accepted 29 august 2017 geological survey of denmark and greenland bulletin 1, 893-930 893 the lower bathonian – middle oxfordian charcot bugt formation is a marginal marine clastic wedge, which has received relatively little attention in the past due to its remote position and its coarse-grained, mostly unfossiliferous nature (bay 1895; aldinger 1935; håkansson et al. 1971; callomon & birkelund 1980). the aims of this study were to combine biostratigraphic and sedimentological data into a coherent depositional and sequence stratigraphic model for the charcot bugt formation in milne land. in previous studies, the biostratigraphy and sedimentology of a basement-onlapping shallow marine sandstone succession, the charcot bugt formation, middle–upper jurassic, east greenland michael larsen, stefan piasecki and finn surlyk a rocky shore developed in early middle jurassic times by transgression of the crystalline basement in milne land at the western margin of the east greenland rift basin. the basement is onlapped by shallow marine sandstones of the charcot bugt formation, locally with a thin fluvial unit at the base. the topography of the onlap surface suggests that a relative sea-level rise of at least 300 m took place in early bathonian – middle oxfordian times. the sea-level rise was punctuated by relative stillstands and falls during which progradation of the shoreline took place. palynological data tied to the boreal ammonite stratigraphy have greatly improved time resolution within the charcot bugt formation, and the jurassic succession in milne land can now be understood in terms of genetically-related depositional systems with a proximal to distal decrease in grain size. the sequence stratigraphic interpretation suggests that translation of the depositional systems governed by relative sea-level changes resulted in stacking of sandstone-dominated falling stage deposits in the eastern, basinwards parts of milne land, whereas thick, remarkably coarsegrained transgressive systems tract deposits formed along the western basin margin. the bulk of the charcot bugt formation consists of stacked sandstone-dominated shoreface units that prograded during highstands. the overall aggradational to backstepping stacking pattern recognised in the charcot bugt formation is comparable to that in the contemporaneous pelion formation of the jameson land basin and in correlative units of the mid-norway shelf and the northern north sea. we suggest that the long-term evolution of the depositional systems may have been controlled by long-term eustatic rise acting in concert with relative sea-level changes reflecting regionally contemporaneous phases of rift initiation, climax and gradual cessation of rifting. keywords: east greenland, milne land, bathonian–oxfordian, charcot bugt formation, kap leslie formation, sedimentology, biostratigraphy, dinoflagellates, sequence stratigraphy, shallow marine, basement onlap, clinoform unit m.l. & s.p., geological survey of denmark and greenland, geocenter copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: mil@geus.dk f.s., geological institute, university of copenhagen, geocenter copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. geological survey of denmark and greenland bulletin 1, 893–930 (2003) © geus, 2003 stratigraphy was based solely on macrofossils; herein we present new correlations combining the existing ammonite stratigraphy with dinoflagellate data. these new biostratigraphic data also allow a better subdivision of the coarse-grained marginal marine sandstones. rocky shorelines are spectacular but rarely described features in the rock record. the onlap contact between the former subaerially exposed rock surface and the marine strata allows quantification of the relative sealevel changes during the middle jurassic. facies analy894 70°45'n 25°30'w 5 km kap leslie kosmocerasdal parat kløft hartz fjeld bays fjelde aldinger elv visdal 25°45'w26°00'w charcot bugt mudderbugt jameson land traill ø scoresby sund 72°n 71°n 26°w 22°w 500 km 25 km g re en la nd charcot gletch er 12 3 4 5 12 13 11 10 9 8 7 6 16 15 19 14 18 17 greenland milne land charcot bugt formation (bathonian – m. oxfordian) kap leslie formation (callovian – m. volgian) hartz fjeld formation (m. volgian – hauterivian) palaeogene basalts quaternary ice crystalline basement section localities fault fig. 1. map showing the distribution of mesozoic sediments in milne land, east greenland. the sandstones of the bathonian – middle oxfordian charcot bugt formation onlap caledonian crystalline basement and are overlain by silty mudstones of the middle callovian – middle volgian kap leslie formation; the boundary between the two formations is diachronous, younging to the west. middle volgian – hauterivian sandstones of the hartz fjeld formation are exposed to the east. the succession is unconformably overlain by palaeogene flood basalts. inset maps indicate the location of milne land, jameson land and traill ø in east greenland; the red line indicates the log panel in fig. 18. sis provides the basis for interpretation of the depositional environments during an overall middle jurassic sea-level rise that resulted in erosion of the former subaerially exposed area and the formation of unusually thick transgressive deposits. the giant-scale cross-sets that form the top of the charcot bugt formation in visdal were first described by callomon & birkelund (1980). three-dimensional mapping of these cross-bedded units has revealed a complex internal upbuilding, each unit being composed of clinoform sets. it is suggested that the clinoforms formed by strong progradation of the clastic shoreline, probably in response to pulses of falling relative sea level. geological setting the late palaeozoic – mesozoic east greenland rift basin is part of the larger rift complex between greenland and the baltic shield that existed prior to the opening of the north atlantic (ziegler 1988; doré 1991). rift basin formation was initiated in devonian times, and late permian through mesozoic basin evolution was governed by cooling and thermal contraction punctuated by phases of extensional faulting, resulting in the development of basin margin half-grabens (surlyk et al. 1984, 1986; surlyk 1990, 1991, 2003, this volume). the east greenland continental margin was uplifted in neogene times and now presents excellent exposures of a mesozoic sedimentary succession deposited in an epicratonic rifted seaway. during the middle jurassic, the jameson land basin formed a north–south elongate embayment, c. 140 km wide and more than 400 km long. the lower bathonian – middle oxfordian charcot bugt formation was deposited during a long-term transgression recognised throughout the east greenland basin (surlyk 1990, 1991, 2003, this volume), and a high gradient rocky shore was 895 40 0 30 0 20 0 10 0 outcrop of onlap surface contour of onlap surface (m) 5 km 25°30′w25°45′w 70°45′n kap leslie p cb cb kl b b cb sw ne p p fig. 2. outcrop of the charcot bugt (cb; c. 100 m thick) and kap leslie (kl) formations at visdal, viewed towards the north-west. note the sharp boundary between the formations interpreted to represent a coincident sequence boundary and marine transgressive surface of erosion. the base of the charcot bugt formation is a marine onlap surface and reflects the pre-jurassic topography of the crystalline basement (b). high peaks are formed of palaeogene flood basalts (p). fig. 3. contour map of the crystalline basement – sediment boundary representing the topography of the onlap surface in bathonian–oxfordian times in milne land. the map is corrected for post-jurassic regional tectonic dip of 3.5° towards the southeast. formed along the faulted western basin margin (fig. 1). in milne land, shallow marine sediments of the charcot bugt formation onlap the irregular, south-east dipping surface of the caledonian crystalline basement (fig. 2; larsen 1995). the orientation of the onlap surface has been corrected for post-jurassic tilting of 3.5° and can be shown to have risen c. 300 m over a distance of 8–15 km in an up-dip direction (fig. 3). this corresponds to an average dip of the basement surface of 1–2°, but locally around basement highs the dip is up to 36°. the amount of relative sea-level rise is estimated by tracing the surface representing the contact between the former subaerially exposed basement and the onlapping marine deposits. stratigraphy the main emphasis of previous studies of the mesozoic of milne land was on the general stratigraphy (bay 1895; rosenkrantz 1929; aldinger 1935), late jurassic ammonite biostratigraphy and lithostratigraphy (spath 1935, 1936; callomon & birkelund 1980; birkelund et al. 1984; birkelund & callomon 1985) and palynology (piasecki 1979, 1980). lithostratigraphy the 1 km thick jurassic – lower cretaceous succession of milne land is subdivided into three formations (callomon & birkelund 1980; birkelund et al. 1984). the lower bathonian – middle oxfordian charcot bugt formation forms the basal sedimentary unit, and consists of coarse-grained sandstones and conglomerates (fig. 4). it is overlain by, and passes laterally into, mudstones and fine-grained sandstones of the callovian – middle volgian kap leslie formation (fig. 4). in the eastern part of milne land, shallow marine and deltaic sandstones of the middle volgian – hauterivian hartz fjeld formation overlie the kap leslie mudstones (piasecki 1979, 1980; birkelund et al. 1984; surlyk et al. 1993). palaeogene flood basalts unconformably overlie the mesozoic succession. charcot bugt formation in outcrop, the charcot bugt formation reaches a maximum thickness of 195 m in the eastern part of milne land and thins towards the west. at visdal (fig. 1), the formation varies in thickness from 162 m in the south to 97 m in the north. the topographic relief of the underlying crystalline basement (fig. 3) controls the general wedge-shaped geometry and local variations in thickness. the occurrence of progressively younger ammonites in the sediments immediately above the unconformity surface towards the west also reflects the onlapping nature of the formation (fig. 4). the upper formation boundary is a strongly diachronous surface which youngs towards the west from the early callovian to the late oxfordian, as indicated by the ammonite and dinoflagellate stratigraphy (fig. 4). the formation is subdivided into the visdal and mudderbugt members (callomon & birkelund 1980). the visdal member forms the main part of the formation and consists of sandy conglomerates with clasts of local crystalline basement rocks overlain by mediumto coarse-grained sandstones with minor siltstone beds. the overlying mudderbugt member forms a south-eastwards thickening wedge of coarse-grained sandstones, up to 6 m thick, exposed only in the southern part of visdal (fig. 1, locality 6). it overlies the visdal member with a sharp, and in places erosional, lower boundary. kap leslie formation the lower callovian – middle volgian kap leslie formation is subdivided into eight members (callomon & birkelund 1980; birkelund et al. 1984). the kosmocerasdal member (lower callovian – upper oxfordian) is the lowest and consists of bioturbated sandy siltstones and fine-grained sandstones; it forms the offshore fine-grained correlative of the upper charcot bugt formation (fig. 4). the member has a maximum thickness of 170 m along the east coast of milne land and thins westwards (callomon & birkelund 1980). in the southern part of visdal, it is only about 10 m thick, although the exact position of the upper boundary in this area is difficult to establish (callomon & birkelund 1980; piasecki 1980). it is overlain by the upper oxfordian aldinger elv member, which forms a wedgeshaped sandstone body thinning westwards from about 90 m on the east coast of milne land to a few metres at bays fjelde (fig. 4; fürsich & heinberg 1983). the aldinger elv member is overlain by dark silty mudstones with abundant glauconitic levels, referred to the bays elv member (upper oxfordian – lower kimmeridgian). the remainder of the kap leslie formation was discussed thoroughly by birkelund et al. (1984) and is not considered further here. 896 ammonite stratigraphy correlation of the boreal lower middle jurassic ammonite zonation of east greenland with the tethyan european ammonite zonation is hindered by faunal provincialism; the faunal horizons in the basal charcot bugt formation are referred to the boreal zonation (callomon 1959, 1972, 1993, 2003, this volume). in this study, two ammonite zones not previously recognised in milne land are reported. the oldest of these is the lower bathonian a. arcticus chronozone, which is represented by a single specimen of arctocephalites cf. arcticus (whitfield) found in the basal beds of the charcot bugt formation at visdal (figs 1, 4, locality 8). this 897 kosmocerasdal member c ha rc ot b ug t fo rm at io n k ap l es lie f or m at io n boreal chronozones chronostratigraphy ba jo ci an ba th on ia n c al lo vi an o xf or di an m m m u u u u l l l l m id dl e ju ra ss ic u pp er ju ra ss ic a. rosenkrantzi a. regulare a. serratum a. glosense c. tenuiserratum c. densiplicatum c. cordatum q. mariae q. lamberti p. athleta e. coronatum k. jason s. calloviense p. koenigi c. nordenskjoeldi c. apertum c. calyx c. variabile a. cranocephaloide a. ishmae a. greenlandicus a. arcticus c. pompeckji c. indistinctus c. borealis caledonian crystalline basement bays elv member west east 12 locality m13 m12 m7 m1 m1 m13 m12 m10 m9 m6 m8 m5 m4 m3 m2 m1 6 5 4 3 m11 9 36 2 18/5 tst coarse-grained sandstone dinoflagellate assemblage, this study tst tst tst hst hst visdal member hst hst hst ? ? fsstcb4 cb3 fsst fsst fsst cb2 cb1 hst hst aldinger elv member 8 7 b a 1 1 2 4 8 2 kmhighstand systems tracthst transgressive systems tracttst falling stage systems tractfsst fine-grained sandstone a ammonite, this study ammonite faunal horizon (callomon & birkelund 1980) mudstone 7 mudderbugt member fig. 4. lithostratigraphy and chronostratigraphy of the charcot bugt formation and lower kap leslie formation; the boundary between these two formations is emphasised by a red line. the formations are interpreted to form genetically-linked depositional systems with the kosmocerasdal member as the fine-grained offshore correlative of the coarse-grained charcot bugt formation. note the westwards backstepping of the charcot bugt formation reflecting the overall transgression during the bathonian – middle oxfordian. cb1–cb4, clinoform units within the charcot bugt formation. ammonite zonation based on callomon (1993). specimen dates the timing of inundation of the milne land area to the early bathonian. however, unfossiliferous sediments exposed in the eastern part of milne land (fig. 1, localities 1, 2) are probably even older, given the progressive westwards onlap onto the easterly dipping crystalline basement surface. the second ammonite zone not previously recognised is the middle bathonian a. ishmae chronozone which is represented by a specimen of arcticoceras harlandi (rawson) found between localities 8 and 9 (figs 1, 4). the upper bathonian a. cranocephaloide chronozone is represented by specimens of arcticoceras/ cadoceras sp.nov.? aff. variabile (spath) and kepplerites tychonis (ravn) from a conglomerate immediately above the contact to the crystalline basement surface at locality 14 (fig. 1; håkansson et al. 1971; callomon & birkelund 1980, fauna m1). the a. cranocephaloide chronozone is probably also represented at locality 1 in the western part of the area, c. 130 m above the basement (fig. 1; callomon & birkelund 1980, fig. 2). at visdal, a succession of medium-grained unfossiliferous sandstones, approximately 60 m thick, overlies the bed containing fauna m1 (a. cranocephaloide chronozone); the c. variabile, c. calyx, c. apertum, c. nordenskjoeldi and p. koenigi chronozones have not been documented by ammonites in milne land. in the uppermost part of the charcot bugt formation in visdal (the mudderbugt member), specimens of perisphinctes (arisphinctes) cf. or aff. maximum (young & bird) indicate the middle oxfordian c. densiplicatum chronozone, probably the c. maltonense subzone (callomon 1961; callomon & birkelund 1980, fauna m7). a complete middle–upper jurassic ammonite succession from the upper callovian p. athleta chronozone to the middle volgian l. groenlandicus chronozone is present in the kap leslie formation at locality 1 and eastwards (fig. 1), and forms a standard for the boreal ammonite zonation (spath 1935, 1936; callomon & birkelund 1980; birkelund et al. 1984). the kosmocerasdal member yields ammonites of the upper callovian p. athleta chronozone (faunal horizon m2 of callomon & birkelund 1980) from concretionary levels 12 m and 30 m above the top of the charcot bugt formation (figs 1, 4, locality 1). ammonites representing the lower oxfordian q. mariae and c. cordatum chronozones and the middle oxfordian c. densiplicatum chronozone are found higher in the succession (faunal horizons m3–m6 of callomon & birkelund 1980). the middle oxfordian c. tenuiserratum and the upper oxfordian a. glosense chronozones are represented in the upper part of the kosmocerasdal member (faunal horizons m8–m10 of callomon & birkelund 1980). the aldinger elv member contains fauna m11 of the upper oxfordian a. glosense and a. serratum chronozones (birkelund & callomon 1980). the base of the bays elv member corresponds to the upper oxfordian a. regulare chronozone and appears to be isochronous throughout milne land (fig. 4; callomon & birkelund 1980; piasecki 1980). dinoflagellate cyst stratigraphy ongoing studies of the jurassic dinoflagellate cyst stratigraphy in east greenland show that the stratigraphic distribution of most species deviates significantly from the distribution reported from the north sea region. the jurassic sediments of east greenland are therefore dated within the framework of the local dinoflagellate stratigraphy (piasecki 1980; s. piasecki and p. milner, unpublished data) that is correlated with the boreal ammonite zonation (fig. 5). the mostly unfossiliferous coarse-grained sediments of the charcot bugt formation have always represented a stratigraphic problem and have hitherto only yielded four horizons with ammonites. however, dinoflagellate cyst assemblages have been found to occur in thin muddy beds and have greatly improved the subdivision and correlation within the formation. in the partly contemporaneous, offshore marine kap leslie formation, ammonites and dinoflagellate cysts occur more abundantly. the dinoflagellate cyst microflora is of bathonian age in the lower part of charcot bugt formation and of callovian and oxfordian age in the upper part of the formation and in the kap leslie formation. the microflora is subdivided into eight stratigraphic assemblages (appendix 1). assemblages 1–5 and 7 are present in fine-grained beds in the charcot bugt formation whereas assemblages 6 and 8 are from silty mudstones of the kap leslie formation (fig. 4). the dinoflagellate cyst zonation shows that the basal unit of the kap leslie formation, the kosmocerasdal member in kosmocerasdal to the east, is of early callovian s. calloviense chron age (assemblage 6). this is considerably older than hitherto believed on the basis of ammonites of late callovian age found 12 m above the formation boundary (callomon & birkelund 1980). samples from fine-grained levels in the upper part of the charcot bugt formation show the presence of sediments of the lower oxfordian c. cordatum, and the middle oxfordian c. densiplicatum chronozones (assem898 899 boreal chronozones abundant occurrence continuous occurrence acme uncertain occurrence al do rfi a al do rfe ns is si rm io di ni um g ro ss ii ka llo sp ha er id iu m h yp or na tu m pa ra go ny au la cy st a sp . ct en id od in iu m s p. cr us so lia p er ire tic ul at a va le ns ie lla d ict yd ia g on ya ul ac ys ta p ec tin ig er a ka llo sp ha er id iu m s p. pa ra go ny au la cy st a re tip ha gm at a ch yt ro ei sp ha er id iu m c hy tro ei de s ch yt ro ei sp ha er id iu m h ya lin a el lip so id ict yu m c in ct um ch yt ro ei sp ha er id iu m c er as te s at op od in iu m p ol yg on al is lit ho di ni a sp on gi os a pa re od in ia p ro lo ng at a m en di co di ni um g ro en la nd icu m w an ae a di gi ta ta w an ae a th ys an ot a w an ae a fim br ia ta ri ga ud el ta a em ul a sc rin io di ni um c ry st al lin um ka lyp te a st eg as ta am bo no sp ha er a ca llo via na d in go di ni um ju ra ss icu m st ep ha ne lyt ro n sp p. ep ip lo sp ha er a bi re tic ul at a lie sb er gi a sc ar bu rg he ns is ev an sia ja ne ae ba jo ci an ba th on ia n c al lo vi an o xf or di an m id dl e ju ra ss ic u pp er ju ra ss ic stratigraphic occurrence of selected species in east greenland a. rosenkrantzi a. regulare a. serratum a. glosense c. tenuiserratum c. densiplicatum c. cordatum q. mariae q. lamberti p. athleta e. coronatum k. jason s. calloviense p. koenigi c. nordenskjoeldi c. apertum c. calyx c. variabile a. cranocephaloide a. ishmae a. greenlandicus a. arcticus c. pompeckji c. indistinctus c. borealis fig. 5. stratigraphic occurrence of selected dinoflagellate cyst species in jameson land, used for dating the assemblages in milne land. 900 f ac ie s li th ol og y st ru ct ur es lo w er bo un da ry bo dy a nd tr ac e fo ss ils pr oc es s in te rp re ta tio n d ep os iti on al en vi ro nm en t 7 ) pe bb ly la g pe bb ly s an ds to ne s m as si ve , l oc al ly t ro ug h cr os sbe dd in g, ra re ri pp le fo rm se ts o n to p su rf ac e sh ar p er os io na l be le m ni te s, am m on ite s, bi va lv es , d ip lo cr at er io n ha bi ch i, s ko lit ho s is p. , m on oc ra te rio n te nt ac ul at um u pp er p ar t of lo w er flo w r eg im e. w av e w in no w in g an d la g fo rm at io n u pp er s ho re fa ce , tr an sg re ss iv e la g de po si ts 6 ) pa ra lle l-l am in at ed sa nd st on e fi ne to m ed iu m -g ra in ed sa nd st on es , l oc al ly m ic ac eo us lo w a ng le in cl in ed pa ra lle l l am in at io n, in te rn al t ru nc at io ns er os io na l u pp er fl ow r eg im e sw as h– ba ck w as h fo re sh or e, b ea ch 5 ) w av e ri pp le cr os sla m in at ed sa nd st on e fi ne -g ra in ed s an ds to ne s lo ca lly e nr ic he d in h ea vy m in er al s c ro ss -la m in at io n, ra re r ip pl e fo rm se ts g ra da tio na l cu rv ol ith os m ul tip le x, pl an ol ite s is p. , sk ol ith os is p. , m on oc ra te rio n te nt ac ul at um fa irw ea th er w av es w av edo m in at ed up pe r sh or ef ac e 4 ) tr ou gh c ro ss be dd ed s an ds to ne fi ne to c oa rs egr ai ne d lo ca lly p eb bl y sa nd st on es tr ou gh c ro ss -b ed di ng , se t he ig ht s 5– 30 c m er os io na l r ar e am m on ite s sk ol ith os is p. , ar en ic ol ite s is p. , es ca pe b ur ro w s w av ege ne ra te d cu rr en ts , m ig ra tio n of t hr ee di m en si on al d un es w av edo m in at ed up pe r sh or ef ac e 3 ) bi oc la st ic co ng lo m er at e pe bb le to c ob bl esi ze d cl as ts o f c or al s an d cr ys ta lli ne r oc ks in a m ed iu m to c oa rs egr ai ne d sa nd st on e m at ri x m as si ve , n or m al ly or r ev er se -t ono rm al ly gr ad ed sh ar p, lo ca lly er os io na l a llo ch th on ou s co ra ls , oy st er s, bi va lv es , be le m ni te s, ra re a m m on ite s su ba qu at ic g ra vi ty fl ow s st or m -in flu en ce d ro ck y sh or ef ac e 1 ) r eg ol ith k ao lin is ed m ig m at ite , pe bb ly r eg ol ith m as si ve g ra da tio na l t o un w ea th er ed cr ys ta lli ne r oc ks su ba er ia l w ea th er in g su ba er ia lly e xp os ed su rf ac e in a w ar m , hu m id c lim at e 2 ) pe bb ly s an ds to ne pe bb ly, m ed iu m to co ar se -g ra in ed fi ni ng up w ar ds in to fi ne -g ra in ed sa nd st on es . c ar bo na ce ou s de br is an d m ic a m as si ve , t ro ug h cr os sbe dd in g gr ad in g in to pa ra lle l l am in at io n sh ar p, e ro si on al r ar e ar th ro po d gr az in g tr ac ks m ig ra tio n of t hr ee di m en si on al d un es . sc ou r an d fil l a llu vi al p la in , ch an ne l f ill t ab le 1 . c ha rc ot b ug t fo rm at io n fa ci es c la ss ifi ca tio n 901 13 ) o ffl ap pi ng cl in of or m s c oa rs e t o ve ry -c oa rs egr ai ne d sa nd st on es fa nsh ap ed c lin of or m se t (1 2 m ) w ith m as si ve or t ro ug h cr os sbe dd ed in tr as et s er os io na l i n up di p pa rt , st ee p do w nl ap in do w ndi p di re ct io n t hr ee -d im en si on al d un es . tr un ca tio n an d st ro ng pr og ra da tio n d el ta o r sh el f pr og ra di ng w ed ge , er os io na l p ha se 12 ) a gg ra di ng cl in of or m s m ed iu m to c oa rs egr ai ne d sa nd st on es , co m m on m ud d ra pe s an d ca rb on ac eo us de br is c om po un d gi an tsc al e cr os sse ts ( 20 m ) w ith lo w -a ng le m as te r be dd in g. pl an ar c ro ss -b ed de d in tr as et s g ra da tio na l, ve ry lo w -a ng le t an ge nt ia l do w nl ap sk ol ith os is p. , p la no lit es is p. tw odi m en si on al d un es m ig ra tin g on c lin of or m su rf ac es . s lo w p ro gr ad at io n an d ag gr ad at io n d el ta o r sh el f pr og ra di ng w ed ge 11 ) pr og ra di ng cl in of or m s m ed iu m to c oa rs egr ai ne d lo ca lly p eb bl y sa nd st on es c om po un d gi an tsc al e cr os sse ts ( 29 m ), w ith hi gh a ng le m as te r be dd in g. tr ou gh c ro ss -b ed de d in tr as et s h ig han gl e do w nl ap r ar e sk ol ith os is p. t hr ee -d im en si on al d un es m ig ra tin g do w n di p on cl in of or m s ur fa ce s. h ig h ra te o f p ro gr ad at io n d el ta o r sh el f pr og ra di ng w ed ge 10 ) t id al ly b un dl ed pl an ar c ro ss -b ed de d sa nd st on es m ed iu m to c oa rs egr ai ne d sa nd st on es w ith a bu nd an t ca rb on ac eo us m ud de br is t id al ly b un dl ed p la na r cr os s be dd in g (u p to 1 .5 m ). ty pe b a nd c r ea ct iv at io n su rf ac es er os io na l st ro ng b io tu rb at io n, sk ol ith os is p. u pp er p ar t of lo w er flo w r eg im e, t w odi m en si on al d un es , eb bdo m in at ed t id al cu rr en ts eb bdo m in at ed tid al c ha nn el 9 ) la rg esc al e cr os sbe dd ed sa nd st on e m ed iu m to c oa rs egr ai ne d sa nd st on es la rg esc al e pl an ar a nd el on ga te t ro ug h cr os sbe dd in g, se t th ic kn es s 0. 5– 4 m er os io na l s et bo un da ri es sk ol ith os is p. , a re ni co lit es is p. u pp er p ar t of lo w er flo w r eg im e, s tr on g tr ac tio n cu rr en ts sh or ef ac e, d un e fie ld s in flu en ce d by lo ng sh or e cu rr en ts 8 ) la m in at ed m ud st on e si lty m ud st on es a nd m ic ac eo us v er y fin egr ai ne d sa nd st on es h or iz on ta l l am in at io n, ra re w av e ri pp le s, ra re h um m oc ky cr os sst ra tif ic at io n sh ar p, p la na r a m m on ite s, bi va lv es ch on dr ite s is p. , p la no lit es is p. , ta en id iu m s er pe nt in um , ra re s ko lit ho s is p. su sp en si on fa ll ou t, ep is od ic r ew or ki ng by w av e an d co m bi ne d cu rr en ts o ffs ho re m ar in e, tr an si tio na l t o st or m -in flu en ce d lo w er s ho re fa ce 902 lo ca lit y 14 n or th so ut h lo ca lit y 15 lo ca lit y 8 lo ca lit y 17 0246 rocky shore association 8 0246 alluvial/shoreface? association 8 0246 shoreface association 8 0 f m ud sa ndm c g r 246 alluvial association shoreface association 8 f m ud sa ndm c g r f m ud sa ndm c g r f m ud sa ndm c g r m m m m k 20 0 0 0 1 2 km m lo c. 1 4 lo c. 1 5 lo c. 8 lo c. 1 7 k ap l es lie f or m at io n c ha rc ot b ug t fo rm at io n fi g. 6 . n o rt h –s o u th g eo lo gi ca l p ro fi le s h o w in g th e c h ar co t b u gt a n d k ap l es lie f o rm at io n s in v is d al , a n d r ep re se n ta tiv e se d im en ta ry s ec tio n s o f th e b as al c h ar co t b u gt f o rm at io n . n o te t h e th ic kn es s va ri at io n o f th e sa n d st o n es d u e to t h e to p o gr ap h y o f th e u n d er ly in g cr ys ta lli n e b as em en t su rf ac e. t h re e su cc es si ve a m m o n ite -b ea ri n g h o ri zo n s ar e in d ic at ed , ill u st ra tin g th e p ro gr es si ve o n la p o f th e ir re gu la r su rf ac e. t h e ac co m p an yi n g le ge n d a ls o a p p lie s to f ig . 8. blage 7) and thus preclude the existence of a major hiatus between the visdal and mudderbugt members, contrary to the interpretation of callomon & birkelund (1980). the dinoflagellate cyst samples from the kap leslie formation in visdal to the west indicate the presence of the upper oxfordian a. glosense and a. serratum chronozones (assemblage 8) in the mudstones directly above the charcot bugt formation. the palynological data allow correlation of the upper sandstone units with the offshore marine mudstones, which were previously interpreted to overlie the sandstones with a major hiatus (callomon & birkelund 1980). the new data thus indicate a genetic relationship between the charcot bugt formation and the kosmocerasdal member of the kap leslie formation (fig. 4). sedimentology thirteen facies are recognised in the charcot bugt formation (table 1). they are grouped into the alluvial, rocky shoreline, shoreface, prograding wedge, and offshore facies associations, each characterised by their constituent facies, geometrical arrangement of facies types, nature of bounding surfaces, and overall geometry. the facies associations are interpreted in terms of depositional environments and correspond to depositional systems. alluvial facies association (facies 1, 2) the alluvial facies association is restricted to the basal part of the formation, mainly situated in local topographical basement lows (fig. 6). the association consists of facies 1 (kaolinised migmatite/regolith) overlain by facies 2 (cross-bedded sandstones) showing an overall fining-upwards trend (table 1). description the crystalline basement consists of gneissic migmatite and banded coarse-grained granitic migmatite (buchernurminen 1979; henriksen & higgins 1988). the rocks are progressively weathered upwards towards the contact with the sediments forming an up to 2.5 m thick regolith (facies 1; fig. 6, locality 17). the overlying massive sandstones (facies 2) are coarse-grained, locally pebbly, arkoses with lenses of subrounded quartz pebbles, lithic fragments and kaolinised and fresh feldspar. 903 m ud st on e (c ar bo na ce ou s) sa nd st on e li th ol og y pe bb ly s an ds to ne c on gl om er at e m at ri xsu pp or te d br ec ci a c ry st al lin e ba se m en t, ka ol in is ed c ry st al lin e ba se m en t pe bb le la g w ea k m od er at e st ro ng pl an ar c ro ss -b ed di ng c lin of or m b ed w ith in tr as et s d ir ec tio n of p al ae oc ur re nt fr om r ip pl es d ir ec tio n of m ig ra tio n fr om cl in of or m d ip bo un di ng s ur fa ce s d eg re e of b io tu rb at io n se qu en ce b ou nd ar y m ar in e flo od in g su rf ac e tr an sg re ss iv e su rf ac e of e ro si on k sb t se fs h el m in th op sis m ag na pl an ol ite s is p. tr ac e fo ss ils cu rv ol ith os m ul tip le x ta en id iu m s er pe nt in um m on oc ra te rio n te nt ac ul at um d ip lo cr at er io n ha bi ch i ro ss el ia is p. sk ol ith os is p. st ru ct ur el es s pl an ar b ed di ng se di m en ta ry s tr uc tu re s pl an ar la m in at io n w av y be dd in g w av e ri pp le c ro ss -la m in at io n tr ou gh c ro ss -b ed di ng be le m ni te a m m on ite o ys te r pl an t fr ag m en t c oa lif ie d w oo d c or al en al lo co en ia c al lo m on i fo ss ils clasts are rounded to well-rounded, spherical and subdiscoidal pebbles of vein quartz, up to 9 cm in diameter. the sandstones grade up into cosets of trough cross-bedded, mediumto coarse-grained sandstone, with pebbly lenses and scattered pebbles (fig. 6, localities 15, 17). foreset azimuth orientations indicate transport directions towards the south-east. even, parallel laminated well-sorted, very fineto fine-grained sandstones form the top of the fining-upwards successions. the lamination is defined by abundant flakes of mica and carbonaceous detritus. bioturbation is restricted to scattered meandering grazing traces (pascichnia) preserved as shallow epireliefs on bedding planes in the thin-bedded upper part of fining-upwards successions. interpretation weathering of the crystalline basement surface and formation of the regolith took place during a period of subaerial exposure. the dominance of kaolinite in the weathering profile may suggest a warm, humid climate (curtis 1990; retallack 1990). the overlying poorly sorted pebbly sandstones, containing reworked subangular lithic fragments, kaolinitic feldspars and quartz, are interpreted to have been deposited in an alluvial environment where clasts were transported only relatively short distances before deposition. the dominance of trough cross-sets and the unimodal palaeocurrent direction perpendicular to the reconstructed palaeo-gradient further support this interpretation (fig. 3; larsen 1995). the sandstones were deposited by three-dimensional dunes in the upper lower flow regime, or represent the fill of shallow scours, and probably formed in shallow channels dominated by bedload transport (miall 1977, 1978; rust 1978, kleinspehn et al. 1984). deposits of the alluvial facies association may have formed a more extensive and continuous cover of the basement surface prior to the middle jurassic transgression, during which the main part was reworked and redistributed by marine currents or waves. the variation in thickness of the regolith may reflect primary variation in depth of weathering or later erosion. rocky shore association (facies 3–6) the rocky shore association consists of bioclastic conglomerates (facies 3), trough cross-bedded sandstones (facies 4), wave ripple cross-laminated sandstones (facies 5) and parallel-laminated sandstones (facies 6; table 1). the facies occur in random successions with abundant internal scour and erosion surfaces. the association is present in the immediate vicinity of crystalline basement highs (fig. 6, locality 14). description the association onlaps the steeply rising crystalline basement surface in the northern part of visdal (fig. 6). in contrast to the low relief basement surface in the southern part of visdal (see above), the crystalline rocks at the onlap surface are generally fresh. neptunian dykes occur where small fractures in the surface are filled with coarse-grained sandstones and well-rounded crystalline pebbles or fragmented oyster shells (fig. 7a). despite a thorough search, no in situ epifauna was found on the basement surface. a well-rounded gneiss boulder more than 1 m in diameter, surrounded by coarse-grained sandstones, occurs in the basal part of the succession near the basement surface at locality 14 (fig. 6). the conglomerates (facies 3) are matrix-supported with pebbleto cobble-sized clasts in a coarse-grained sandy matrix. clasts may be polymict, comprising quartz pebbles, lithic fragments, belemnites, fragments of corals and oyster shells (fig. 7b) or they may be composed exclusively of fragments of the coral enallocoenia callomoni (beauvais) (fig. 7c). the corals have been affected by bioerosion and gastrochaenolites borings, representing resting nests of bivalves such as lithophaga or gastrochaena, are common (fig. 7d; frey & seilacher 1980). the conglomerates form tabular beds, up to 50 cm thick, with a sharp, locally erosional base and inverse-to-normal or normal grading. well-sorted fine-grained sandstone caps, up to c. 30 cm thick, show parallel lamination. bedding surfaces locally show horizontal traces of taenidium serpentinum and helminthopsis magna. the conglomerates are associated with fineto coarse-grained sandstones that are pebbly in places (facies 4). the sandstones are trough cross-bedded with set thickness between 5 and 30 cm. high angle scours may be partly filled with pebbly sandstones, and clast-supported conglomerate sheets occur interbedded with the cross-sets (fig. 7e). foreset orientations in the trough cross-beds indicate dominant transport towards the south. the coarse-grained facies are interbedded with wellsorted fineto medium-grained wave ripple cross-laminated sandstones (facies 5), showing intense burrowing 904 905 fig. 7. examples of sedimentary facies and faunas of the rocky shoreline association. scale in centimetres. a: subvertical contact between a coquina and unweathered crystalline basement. the shelly debris was deposited in a fracture in the basement surface forming a neptunian dyke. b: polymict siliciclastic and bioclastic conglomerate with lithic fragments, quartz pebbles, fragmented corals (c), oysters (o), and worn well-rounded belemnites (b). c: marine matrix-supported conglomerates with coral clasts. photograph shows vertical stacking of three graded beds. bed boundaries indicated by dashed lines. d: compact hemispherical colony of the hermatypic coral enallocoenia callomoni beauvais recovered from a conglomerate. the coral shows sand-filled gastrochaenolites borings probably representing nests of boring bivalves. e: pebbly sandstones filling a shallow scour interpreted as a rip-channel fill (lower part, under scale). note the gradual upwards decrease in the dip of the foresets and concomitant decrease in grain size. the rip-channel fill is overlain by parallel-laminated sandstone (above scale). f: wave-rippled sandstones truncated by pebbly trough cross-bedded sandstones. note the abrupt changes in grain size and the abundant erosional surfaces. a e f b b o c c dc 2 cm 906 by curvolithos multiplex, and parallel-laminated wellsorted fineto medium-grained sandstones, locally rich in mica (facies 6). bed thicknesses reach 30 cm, but may be reduced due to truncation by pebble-filled scours (facies 4; fig. 7f). the parallel lamination of facies 6 is defined by coarse sand grains or fine quartz pebbles and small coralline fragments. the lamination dips a few degrees and shows low-angle internal truncations. interpretation the generally poor sorting of the pebbly sandstones and the composition of the conglomerates point towards mixing of marine and terrestial sediments, suggesting marine reworking of alluvial sediments. textural sorting and marked grain-size segregation between individual sets are characteristic of sediments deposited in the ambient energy regime on a wave-dominated upper shoreface (dupré et al. 1980; clifton 1981; nemec & steel 1984; leithold & bourgeois 1984). the trough cross-stratified pebbly sandstones were deposited by strong unidirectional currents that may have been generated by storm surges either along the coast or in rip channels. the matrix-supported graded conglomerates contain dispersed floating clasts, a fabric that invites comparison with debris flow deposits. the out-sized clasts, however, consist primarily of corals, which may have shown a different hydrodynamic behaviour from siliciclastic material of equal size. the floating clasts therefore are not unequivocal criteria with respect to the depositional process. however, the overall inverse-to-normal grading present in some beds and the lack of traction current generated structures support the interpretation that they represent high density gravity flow deposits (lowe 1979, 1982). the cross-laminated fine-grained sandstones (facies 5) that form a cap on some of the conglomerate beds are interpreted to represent reworking by fair-weather waves. the shallow water depth in front of the rocky shoreline is reflected by the parallellaminated sandstones formed by high-energy swash– backwash on the foreshore or by shoaling waves on the upper shoreface (clifton 1969). the well-rounded gneiss boulder embedded in marine sediments indicates exposure to wave processes probably on a high-energy beach before final burial (dott 1974; surlyk & christensen 1974). it may have originated from wave erosion of the rocky coastline removing a joint-controlled weathering crust leaving the rounded boulder in place or it may have rolled or slid into the site of deposition. the rocky shore association is bounded by the crystalline basement in a landwards direction. the association may be characterised as amalgamated marine lag deposits formed by progressive wave erosion of the basement and alluvial deposits and by winnowing of shoreface and foreshore deposits in a high energy rocky shoreline environment. shoreface association (facies 4, 5, 7–9) stacked units of the shoreface association form the bulk of the charcot bugt formation. each unit shows a basal erosional surface overlain by a coarse-grained lag deposit (facies 7). the lag is sharply overlain by silty mudstones forming the base of an overall coarsening-upwards succession (facies 8, 4, 5, 9), up to 22 m thick (fig. 8). description the basal lag deposits reach a maximum thickness of 1 m and are composed of coarseto very coarse-grained quartzitic sandstone with concentrations of worn belemnites and ammonites. the lags are intensely burrowed by skolithos isp. and diplocraterion habichi (facies 7; fig. 8). the upper bedding surfaces are planar or locally reworked into large-scale wave ripples. the lag deposits are abruptly overlain by laminated mudstones (facies 8) that form the lower part of the coarsening-upwards successions. the siltstones grade upwards into well-sorted mica-rich, very fineto fine-grained sandstones, showing wave ripple cross-lamination and symmetrical ripples on bedding planes (facies 5). the sandstones are strongly bioturbated and show well-preserved curvolithos multiplex and planolites isp. the fine-grained sandstones are overlain by cosets, 6–10 m thick, of trough cross-bedded, medium-grained sandstones (facies 4). individual sets are medium-scale trough cross-beds with set thickness between 5 and 25 cm and low-angle foresets. the mudderbugt member consists solely of coarse-grained sandstones of facies 4. in visdal, the uppermost of the shoreface units contains cosets of very large-scale, planar and trough cross-bedded, mediumto coarse-grained sandstones (facies 9; figs 8, 9). the cosets have a sheet-like geometry and are bounded by erosional planar surfaces. they are up to 9 m thick; individual sets are up to 4 m thick, with 0.5–1.5 m being most common. the set boundaries form more than 10 m wide troughs, with erosional boundaries to underlying sets (fig. 9). foresets are dominantly tangential or more rarely sigmoidal. individual cross-sets show inversely or, rarely, normally graded simple avalanche foresets up to 40 cm thick. locally, the foresets are compound showing internal lowangle trough cross-bedding. the large-scale foresets dip up to 23° with a consistent dip direction towards the south (mean 196°; fig. 10). bioturbation is rare; isolated arenicolites isp. and skolithos isp. burrows occur locally. 907 scree covered fs tse fs fs sb/tse fs tse fs tse fs tse fs tse 60 40 30 20 10 0 50 m 110 100 90 80 70 m 4 4 7 4 8 11 11 7 8 11 8 9 9 7 4 4 4 8 7 8 8 4 5 5 4 7 120 cl si f m cgr sand cl si f m cgr kap leslie formation sh or ef ac e as so ci at io n c ha rc ot b ug t fo rm at io n, v is da l m em be r c ha rc ot b ug t fo rm at io n, v is da l m em be r pr og ra di ng w ed ge a ss oc ia tio n cb4 sand fig. 8. sedimentological section measured through the upper part of the charcot bugt formation showing stacked coarsening-upwards shoreface successions overlain by clinoform unit cb4. the small numbers indicate facies, as described in the text. for legend, see fig. 6. interpretation the lateral continuity of the erosionally-based pebbly sandstones and the presence of marine macrofossils and wave-generated ripples suggest that the lag (facies 7) formed by shoreface ravinement followed by shallow marine winnowing (swift 1968; demarest & kraft 1987; nummedal & swift 1987). the trace fossils belong 908 20 m tse cb4 sb/tse west east shoreface strike and dip direction coarseningupwards offshore offshore prograding wedge fig. 9. prograding clastic wedge (cb4) forming the top of the charcot bugt formation at visdal. the unit is more than 50 m thick and downlaps onto marine mudstones that themselves succeed a coarsening-upwards shoreface succession lowermost in the photograph. view towards the north-east. the line drawing is based on the photograph and measured vertical sections. note the steeply dipping clinoforms in the central part of the photograph. person encircled for scale. 909 to the habichi ichnocoenosis (heinberg & birkelund 1984) and are interpreted to indicate shallow marine to intertidal, high energy environments characterised by highly varying sedimentation and erosion rates (heinberg & birkelund 1984; dam 1990). the overlying finegrained sediments are interpreted as having been deposited from suspension fall-out below storm wave base and the transition thus indicates an abrupt increase in water depth and is designated a marine flooding surface. although this surface is easily recognised in outcrop, the definition of the facies associations as genetically-related facies implies that the lower boundary of cycles in the shoreface association is placed at the erosional base (marine surface of erosion) of the coarse-grained lag (see arnott 1995). the coarsening-upwards part of the successions is interpreted to represent shoreface progradation. the fine-grained strongly bioturbated sandstones indicating slow sedimentation rates and deposition in a low energy environment were deposited on the lower shoreface. the upper shoreface is represented by mediumand coarse-grained trough cross-bedded sandstones (facies 4) deposited by three-dimensional dunes in a highenergy wave-regime. the near-shore shallow marine sand sheets and shoreface environments are characterised by a high degree of reworking (dupré et al. 1980; clifton 1981). the high energy, shallow marine environment is also reflected by the trace fossils which are dominated by vertical burrows probably inhabited by suspension feeders. the interpretation of the very large-scale cross-sets (facies 9) is less straightforward. no unequivocal evidence of wave or tidal action is present and the consistent palaeocurrent direction, indicated by the dip of the very large-scale foresets, suggests unimodal currents. based on the sedimentary structures alone the sandstones may have been deposited by large sandbars or mouthbars in either fluvial, tidal or shallow marine environments. the association with upper shoreface facies and the presence of marine trace fossils (albeit scarce), however, suggest a tidal to shallow marine depositional environment. tidally-driven current systems are often separated into an ebband a flood-dominated thalweg, and the lack of tidal structures may therefore not be significant. berné et al. (1991) described large-scale dunes from recent subtidal environments, that were 0.7–9.4 m high with common heights between 2–4 m and formed in water depths between 8 and 23 m. they migrated under the influence of tidal currents. large-scale dunes may also form on epicontinental platforms by strong geostrophic currents, occasional storm surges and/or tidal currents (mccave 1971; flemming 1978; field et al. 1981). the sandstones (facies 9) are therefore interpreted to have been deposited by fields of linear or slightly sinuous subtidal dunes on the shoreface. the dune fields migrated southwards under the influence of coast-parallel currents. surlyk & noe-nygaard (1991) described cross-bedded sandstones of similar scale from the volgian raukelv formation in central jameson land, interpreted as having been deposited in dune fields driven by coast-parallel tidal currents. prograding wedge association (facies 10–13) four clinoform units (cb1–cb4) characterised by highangle foresets or clinoforms are identified in the charcot bugt formation. the units are bounded by major marine flooding surfaces or erosional surfaces, and consist of facies 10–13 (table 1). they are of ?middle–late bathonian (cb1), late bathonian (cb2), callovian (cb3) and early–middle oxfordian (cb4) age, and form a backstepping succession of sandstone-dominated clinoform units (fig. 4). 10 20 30 40 % n = 19 v = 196° 10 20 30 40% n = 103 v = 126° 10 20 30 40% n = 56 v = 157° shoreface foreset dip prograding wedge (clinoform dip) prograding wedge (intraset) fig. 10. equal-area rose plots showing the palaeocurrent direction in the shoreface association (facies 4, 9) and the clinoform dip direction and intraset palaeocurrent direction in the prograding wedge association (facies 10–13). 910 description the four clinoform units were mapped out in the field and have a bank or wedge-shaped geometry (in the sense of mitchum et al. 1977). the maximum thicknesses of cb1, cb2, cb3 and cb4 are 21 m, 19.5 m, 31 m and 50.5 m, respectively. the four units are characterised by steeply dipping clinoforms extending from the top to the base of the unit (facies 11–13; fig. 9). the clinoforms downlap onto the top of large-scale cross-bedded sandstones (facies 10) or the offshore association. in a basinwards direction, the sandstone-dominated clinoform units pass into, and are overlain by, mudstones of the offshore facies association (the kosmocerasdal member). the clinoforms are tangential or locally sigmoid with an inclination of 5–20° towards the south-east, although the steepest middle part of the clinoforms may locally reach 24° (fig. 9). clinoforms are asymptotic towards the lower boundary where they merge into a carbonaceous fine-grained clinoform toe. the sigmoidal clinoforms gradually merge upwards into a horizontally bedded, wave-influenced top, whereas the tangential oblique clinoforms are erosionally truncated at the top. the clinoforms are grouped into progradational, aggradational and offlapping clinoform sets (sydow & roberts 1994), according to the geometry of the clinoform surfaces (fig. 11). the clinoforms show intrasets of planar cross-strata or cosets of trough cross-strata that are characteristic of the different types of clinoform sets (facies 11–13; table 1). the clinoform sets are bounded by erosional surfaces of local extent and/or their correlative surfaces. progradational clinoform sets (facies 11) consist of steeply dipping (15–24°) clinoforms showing low-angle trough cross-bedded intrasets with set thicknesses up to 8 cm. the facies consists mainly of well-sorted coarsegrained sandstone with a low content of carbonaceous debris. bioturbation is weak. aggradational clinoform sets (facies 12) are sigmoidal or tangential with planar cross-bedded intrasets up to 1.4 m thick. clinoforms dip 5–14° and consist of wellsorted fineto medium-grained, locally coarse-grained sandstones with abundant carbonaceous debris. bioturbation is moderate to strong. offlapping clinoform sets (facies 13) are entirely composed of trough cross-bedded coarse-grained sandstones, forming cosets up to 12 m thick (fig. 11). bioturbation is absent except for diplocraterion habichi extending downwards from the upper clinoform set boundary. offlapping clinoform sets are typically found in the upper part of clinoform units and represent the final progradation of the clastic wedge. they show a downstepping (offlapping) geometric arrangement in csb csb csb c lin of or m s et c p b o t o a cu c m m a sb/tse ds c lin of or m u ni t ds downlap surface sb/tse coalesced sequence boundary and transgressive surface of erosion csb clinoform set boundary c clinoform surface p prograding clinoform set a aggrading clinoform set o offlapping clinoform set m offshore mudstones cu condensed unit b bottomset t topset fig. 11. generalised section (scale arbitrary) through a clinoform unit showing the bounding surfaces, structures and geometric characteristics of the clinoform sets. three types of clinoform sets are described based on the geometry and internal structures. they comprise: prograding (p) (facies 11), aggrading (a) (facies 12) and offlapping (o) (facies 13). the clinoform sets are bounded by erosional surfaces and/or fine-grained condensed units. 911 the direction of migration. up-dip, the offlapping clinoforms may be traced into a strongly erosional surface. clinoform unit cb1 cb1 is exposed in steep cliff faces at kosmocerasdal and parat kløft (fig.1, localities 1, 2). minimum areal extent of the clinoform unit is 3 km2 and it can be followed in continuous outcrop for c. 1 km along strike. cb1 is 21 m thick in parat kløft and 17 m thick in kosmocerasdal and consists of mediumto very coarse-grained sandstones showing weak bioturbation by skolithos-type burrows. it consists of a single set of progradational clinoforms. the clinoforms are tangential with truncated upper parts and show maximum dip angles of 16° and migration direction towards the south-east (132°; fig. 12). the clinoform surfaces bound single sets of trough cross-bedded sandstones. the geometry of cb1 cannot be determined, but it forms a tabular sandstone body in outcrop. it downlaps onto mediumto coarse-grained shallow marine sandstones and has a sharp truncated top. cb1 is overlain by coarse-grained sandstones at parat kløft (locality 2) and dark silty mudstones at kosmocerasdal (locality 1). the mudstones at kosmocerasdal contain marine dinocysts (assemblage 3; appendix 1) and a tentative stratigraphic correlation with the established stratigraphic scheme of jameson land suggests a late bathonian, a. cranocephaloide chron or older age (fig. 4; piasecki 1980). this is supported by an ammonite fragment found in the transgressive lag at the top of cb1 in kosmocerasdal (callomon & birkelund 1980, fig. 2). clinoform unit cb2 cb2 is dominated by an aggradational clinoform set. it is exposed in the area around parat kløft (locality 2) over a minimum area of 3 km2 (fig. 12). the geometry of cb2 cannot be determined, due to limited exposure. the clinoforms are sigmoidal and show dip angles up to 15° with a migration direction towards the east (84°) whereas the intrasets show a migration direction towards the south-east (132°; fig. 12). the sigmoidal clinoforms show a very gradual downlap with a rather thick carbonaceous toeset. clinoforms bound planar cross-sets, up to 1.8 m thick, of mediumto coarse-grained sandstones rich in carbonaceous detritus and with scattered silicified wood. bioturbation is restricted to localised skolithos isp. burrows. cb2 is capped by a wave-influenced unit and is overlain by trough crossbedded, mediumto coarse-grained sandstones with belemnites. the maximum thickness of cb2 is 19.5 m; it is late bathonian in age corresponding to the c. calyx and c. apertum chrons based on dinoflagellate cysts (assemblage 4) found in the fine-grained toeset deposits. clinoform unit cb3 cb3 consists of a single set of progradational clinoforms. the clinoform unit is exposed over 8 km2 at kosmocerasdal (figs 1, 12, locality 1) and parat kløft (locality 2) where it forms the uppermost unit of the charcot bugt formation. clinoforms are tangential to sigmoidal and show dip angles up to 20° with a migration direction towards the east (90°, locality 1) and south-east (136°; figs 1, 12, locality 2). clinoforms bound single sets of trough cross-beds up to 8 cm thick, and each set can be traced down-dip for several metres. the trough cross-bedded sandstones are mediumto coarse-grained with scattered carbonaceous detritus and kaolinitic mudstone clasts. cb3 shows localised bioturbation (skolithos isp.) and belemnites occur in the uppermost part of cb3 in kosmocerasdal (figs 1, 12, locality 1). cb3 downlaps onto horizontally bedded or low-angle cross-bedded, mediumto coarse-grained, shallow marine sandstones (facies 4). at kosmocerasdal, cb3 shows a strongly truncated top and is overlain by a fine-grained sandstone bed, 30 cm thick, containing 70°45′n kap leslie 5 km 25°30′w25°45′w cb4 cb3 cb1,2 50 m 0 m direction of migration (clinoform dip) clinoform units cb1, cb2 clinoform unit cb3 clinoform unit cb4 with sandstone isopach fig. 12. isopach map showing thickness variation and geographical distribution of the four clinoform units cb1–cb4. scattered fine pebbles and abundant coalified wood pieces, whereas it shows a wave-influenced top at parat kløft. the sandstones are overlain by sandy siltstones of the kap leslie formation. the clinoform unit has a wedge-shaped geometry with a planar lower surface and an eastwards inclined upper surface. contours of the upper formation boundary (i.e. the upper boundary of cb3 in this area) show that the surface is planar and dips a few degrees towards the south-east. dinoflagellate cysts in the bottomset (assemblage 5) indicate an early callovian age corresponding to the c. nordenskjoeldi or earliest p. koenigi chrons. dinocysts from the basal beds of the overlying kap leslie formation (assemblage 6) indicate the middle callovian s. calloviense or possibly k. jason chronozones (fig. 4; piasecki 1980). the oldest ammonites present in the overlying mudstone succession are found c. 12 m above cb4 and belong to the upper callovian p. athleta chronozone, k. proniae subchronozone (callomon & birkelund 1980). clinoform unit cb4 clinoform unit cb4 forms the uppermost part of the visdal member in the western outcrop area and is exposed in laterally continuous outcrops between localities 6 and 15 and between localities 4 and 5; it covers at least 40 km2 (figs 12, 13). the clinoforms downlap onto marine carbonaceous mudstones in most of the area, but it erosionally overlies medium-grained sandstones of the shoreface association in the northern part of visdal (fig. 13; localities 11–13). the top of cb4 forms the upper formation boundary and is overlain by silty mudstones of the kap leslie formation (localities 8–13). the mudderbugt member erosionally overlies cb4 in the southern part of visdal (localities 6, 7). the age of cb4 is early–middle oxfordian based on dinoflagellate cysts (assemblage 7) found in the underlying marine mudstones and the carbonaceous toesets that are indicative of the c. cordatum and c. densiplicatum chronozones (fig. 4). the age given by the dinoflagellate cyst assemblage is in accordance with ammonites in the overlying mudderbugt member (c. densiplicatum chronozone; callomon 1961). cb4 shows a lobe-shaped geometry with the maximum thickness in the proximal (western) area (figs 1, 12, localities 8–13). the gradual decrease in thickness towards the east and south-east can be observed at outcrop between localities 5 and 4. around locality 4, thin beds of coarse-grained sandstones are found embedded in offshore mudstones, probably representing the most distal part of the clinoform unit. cross-bedded tidal sandstones of facies 10 occur in the proximal (western) part of clinoform unit cb4 and consist of mediumto coarse-grained, generally wellsorted sandstones with abundant disseminated carbonaceous material. tidally-influenced planar and trough cross-beds up to 1.5 m thick, (typically 0.30–0.45 m) form cosets up to 15 m thick (fig. 14). the cosets generally fine upwards and are capped by fine-grained carbonaceous mudstone layers. the planar cross-bedded sets show rhythmic variation in bundle thickness, double mud drapes, and type b and c reactivation surfaces indicating tidal influence (de mowbray & visser 1984; nio & yang 1991a). double mud drapes, less than a few millimetres thick, extend from the bottomsets to approximately two-thirds up the set. reactivation surfaces dip approximately 23° as compared to the maximum foreset dip of 26°. foreset inclinations change systematically, decreasing with decreasing bundle thickness. palaeocurrent readings from the planar cross-sets indicate a dominant southerly transport direction (mean 176°). small reversed current ripples showing foreset azimuths towards the north climb up the reactivation surfaces and form sets up to 1.5 cm thick, draped by a layer of carbonaceous mudstone. the tidally-influenced planar cross-beds show bioturbation concentrated along set boundaries and on the foreset laminae (fig. 14b). the clinoform sets can be followed between outcrops in the valley of visdal, and are arranged in a shingled fashion, such that superposed sets are offset in a southwards direction. the clinoform sets in cb4 show all three geometrical styles defined above. the progradational clinoform sets are most common in the central part of visdal (localities 8 and 9), where they form the lower part of cb4 (fig. 13). the thick progradational clinoform sets are associated with the maximum thickness of the clinoform unit. the progradational clinoform sets reach a maximum thickness of 29 m at locality 8. the clinoforms are tangential, with angles typically varying between 5° and 20°, although locally up to 24°. the clinoforms bound sandstones showing low-angle trough cross-sets, 3–14 cm thick (fig. 15). the aggradational clinoform sets are most common in the northern part of the visdal valley (localities 11, 12), where the low-angle sigmoidal clinoforms bound sets of planar cross-sets (fig. 16). clinoform dip angles are in the range 3–12°. the internal cross-sets are up to 1.5 m thick. the offlapping clinoform sets occur in the upper half of cb4 and are characteristically massive in appearance. the maximum thickness is 12 m (locality 9) and the offlapping sets are composed of cosets of low-angle trough cross-beds. 912 913 the clinoforms of cb4 show a consistent dip towards the south-east (mean of 123°; fig. 12), indicating progradation in this direction, slightly oblique to the reconstructed palaeoshoreline. the internal cross-bedding is slightly offset to the south of the progradation direction and shows a mean palaeocurrent direction towards 157°. clinoform dip angles decrease asymptotically towards the lower boundary and merge into a clinoform toeset unit, several metres thick, rich in carbonaceous detritus. sigmoidal clinoforms in the upper part merge into a horizontally bedded, wave-influenced clinoform top. it varies in thickness dependent on the amount of truncation and reaches a maximum thickness of 6.6 m at locality 9. in the southern part of the visdal valley (locality 7), the upper part of cb4 is truncated by large-scale concave-up erosional surface overlain by steeply dipping avalanche beds. the scour can be followed along strike for more than 1 km and downdip for approximately 100 m. the dip angle of successive avalanche beds in the scour fill decreases gradually upwards and down-dip from 30° to 6°. the top surface of cb4 shows a characteristic proximal (landwards) to distal (basinwards) development. the surface has been walked out in the visdal area 50 m 40 14 67 30 20 10 0 e mudderbugt mb e sse w 13 5 n 12 11a 11b w 8 9 10 4 fig. 13. fence-diagram of clinoform unit cb4 based on vertical sections at visdal and bays fjelde. the cross-sections are somewhat idealised in order to show the component clinoform sets and the nature of the bounding surfaces. cb4 prograded towards the east and south-east and passes laterally into offshore mudstones (green) of the kap leslie formation. localities are shown in fig. 1. 914 where it forms a spectacular, and easily recognisable surface separating cliff-forming coarse-grained sandstones from soft silty mudstones (fig. 2). in northern visdal, localities 9–15, the surface is planar and roughly horizontal. south of locality 9, however, the surface dips approximately 6° towards the south-east. small remnants of tidally-influenced sandstones erosionally overlie the uppermost prograding wedge association. the preserved thickness of this tidal unit varies along the valley and reaches a maximum of 5.5 m around locality 7. the boundary between cb4 and the tidally-influenced cross-sets is locally draped with fine pebbles. at locality 4, cb4 wedges out and the flooding surface capping cb4 coalesces with a flooding surface forming the top surface of the underlying sandstone unit. in the proximal area (localities 9–13), the upper boundary of cb4 is a strongly erosional surface truncating mediumto coarse-grained sandstones. the sandstones immediately beneath the surface are cemented by siderite and show a characteristic reddish weathering colour in the uppermost 5 cm. the surface shows a microrelief of a few centimetres and is covered by a pebble lag, one clast thick (fig. 17). the pebbles are subrounded to wellrounded, subprismoidal or spherical and consist of quartz and lithic fragments. the long axes are up to 4 cm long at locality 12, 2.5 cm at locality 11, and 3 cm at locality 9. the clasts form a laterally extensive pavement in the proximal area but are more dispersed basinwards. the trace fossils skolithos isp. and monocraterion tentaculatum are common, with long vertical tubes descending from the top surface. 1 m n 20 cm fig. 14. sedimentary facies of the prograding wedge association, locality 11, northern visdal. a: coset of tidally-influenced planar cross-bedded sandstones; set boundaires indicated by dashed lines, coset boundary by the solid line. area outlined shown in detail in fig. 14b. b: close-up of tidal cross-beds. note the variation in bundle thickness interpreted as a cyclic variation controlled by the tidal cycle. neap-tide bundles (n) are draped by carbonaceous mudstones. note burrows extending down from set boundaries and at right-angles to the foreset surfaces. bundle thickness increases and mud-drapes and bioturbation decrease from neap to spring-tide bundles. a b 915 interpretation the deposition of large-scale cross-bedded units in marine settings has been attributed to a number of different depositional settings (see review by pomar & tropeano 2001); a characteristic feature of these deposits is that they comprise coarse-grained laterally extensive bodies encased in fine-grained shelf or offshore deposits. the clinoform units described in this study may be interpreted to have formed by progradation of gilberttype deltaic lobes or progradational wedges deposited at the shoreface edge (transition-slope deposits of pomar & tropeano 2001). as stressed by pomar & tropeano (2001), however, differentiation between large-scale clinoform units formed as shore-parallel tabular bodies and deltaic (gilbert-type) lobes is only possible if extensive 3-d data are available. in this study, 3-d outcrops of the individual clinoform units are present in visdal (fig. 1), but regionally their large-scale geometry is poorly constrained. we therefore adopt the purely descriptive term clinoform unit for the large-scale cross-bedded sandstone beds rather than using a generic term such as transition-slope lithosome or delta lobe. the clinoforms represent the depositional surfaces of the prograding wedges at specific times. progradation was largely controlled by the migration of ripples and dunes transporting sediment across the shelf/delta platform and down the front of the prograding wedge. the migration of these bedforms was at a slightly oblique angle to the main direction of progradation and was probably influenced by southwards-directed longshore currents. wave influence was rather limited and seems fig. 15. a: prograding clinoform set in visdal (cb4, locality 8) showing clinoforms dipping towards the south-east (to the right). the clinoforms bound sets of small-scale low-angle trough cross-sets. scale (encircled) is 20 cm long. b: close-up of a showing clinoforms (arrows) with internal trough cross-sets. the ripples forming the intrasets migrated at a slightly oblique angle to the progradation direction of the clinoforms. scale divisions are centimetres. a b 916 1 m fig. 16. a: prograding clinoform set in visdal (cb4, locality 9) showing clinoforms dipping towards the viewer in a south-easterly direction. the clinoforms bound small-scale low-angle trough cross-sets and planar cross-sets. outlined area shown in fig. 16b. b: close-up of a showing planar cross-bedded intrasets. scale is 20 cm long. fig. 17. pebbly sandstone lag (bedding plane view) deposited on the coalesced sequence boundary and transgressive surface of erosion forming the upper boundary of cb4. locality 12, northern visdal; scale in centimetres. a b only to have influenced deposition in the uppermost part of the clinoform unit, whereas tidal influence was locally strong. the tidal currents were probably restricted to channels and their deposits may be characteristic of certain phases of the progradation event. the base of the prograding wedge (the toe of the clinoforms) was situated below storm wave-base as indicated by the absence of wave-generated structures. the water depth in front of the prograding wedge can be estimated from the thickness of the clinoform unit and was at least 50 m during the progradation of cb4. the basin floor was dominated by deposition of silt and mud from suspension accompanied by strong bioturbation. progradation was not a continuous process, and can be divided into phases based on the different types of clinoform sets (facies 11–13). each of the clinoform types is thus interpreted to represent a particular progradational phase controlled by autocyclic processes (sediment supply) and/or changes in relative sea level. the clinoform units cb1, cb2, and cb4 are dominated by the steep clinoforms of the prograding clinoform sets (facies 11). the rather high angle of the clinoforms is unusual, but such angles have been described as a characteristic feature of bed-load dominated systems prograding during a fall in relative sea level (posamentier & morris 2000). the change in sea level results in erosion of previously deposited sediments and hinders aggradation. this creates a high relief front of the progradational wedge with steeply dipping foresets as the depositional system is forced to prograde into deeper water. the high inclination of the clinoforms is controlled firstly by an abrupt deepening in front of the prograding wedge, secondly by the coarse grain size and thirdly by progradation during falling sea level. examples of such steeply dipping large-scale foresets have been described from the recent alta delta (norway) in which the foresets dip at angles of 8–37° (corner et al. 1990), from the campanian panther tongue member in utah (posamentier & morris 2000), in which the clinoforms of the delta front are up to 15 m high and dip at angles up to 27°, and from the calcarenite di gravina formation in southern italy (pomar & tropeano 2001) which displays large-scale cross-bedding with foreset dips up to 35°. in cb4, the progradational clinoform sets locally show a gradational upwards change to planar cross-bedded low-angle clinoform sets (facies 12; aggradational). this may reflect the build-up of smaller depositional units on the shelf platform in response to short periods of rising sea level. the platform deposits are capped by a finegrained and strongly bioturbated sandstone unit which represents the abandonment of the depositional system in response to rising sea level or a shift in the current system transporting sediment to the prograding wedge. the progradation of the coarse-grained unit probably took place during a short time interval relative to the accumulation of the fine-grained carbonaceous facies. the downstepping of successive clinoform sets (facies 13; offlapping) accompanied by erosion of the previously deposited sediments suggests that deposition occurred in response to falling relative sea level. certain clinoform sets can be followed in outcrop and show a southwards shingling. the southwards migration of the depositional system was probably controlled by southerly-directed currents, which also dominated deposition of the shoreface facies association. in the proximal western area, the very low-angle depositional surfaces and the well-developed planar cross-sets indicate deposition under the influence of tidal currents. the large-scale planar cross-sets, up to 1.5 m thick, reflect deposition by large-scale two-dimensional dunes. the dunes are ebb-dominated as shown by the southerly current direction. the reactivation surfaces, systematic changes in bundle thickness and reversed current ripples associated with double mud drapes testify to active tidal currents during dune migrasion and suggest a strongly asymmetric and rectilinear character of the tidal current ellipses (nio & yang 1991a, b). deposition of mud drapes and formation of reactivation surfaces would be strongly favoured by channelised tidal currents with pronounced slack water periods and a relatively weak wave influence characterising an in-channel depositional environment (nio & yang 1991a, b). the geometry of the channel cannot be recognised but a channel environment is supported by the erosional lower boundary and the fining-upwards and thinning-upwards trends. offshore association (facies 8) the offshore facies association consists of parallel-laminated silty mudstones with thin sandstone layers and concretionary horizons (facies 8). the association forms the bulk of the kosmocerasdal member, representing the distal equivalent to sandstones of the charcot bugt formation (fig. 4). description at kosmocerasdal (fig. 1, locality 1), the association is characterised by silty and sandy mudstones forming 917 coarsening-upwards successions up to 20 m thick, capped by a sandy bed or a concretionary layer with abundant ammonites. the mudstone succession is otherwise unfossiliferous and bioturbation is rare. the total organic carbon (toc) content is below 1% which is characteristic for the lower part (callovian) of the kosmocerasdal member. at visdal, the basal few metres of the offshore succession (middle oxfordian) consist of black finely laminated mudstones with abundant small chondrites isp. burrows. several horizons are glauconitic. above the basal unit, the succession consists of monotonous mudstone units interbedded with thin layers and concretionary horizons of fine-grained sandstones. geochemical analysis shows toc values of 4–7%, decreasing upwards to less than 1%. macrofossils are rare and restricted to ammonites and bivalves. well-sorted, erosionally based, fine-grained sandstones, up to c. 30 cm thick, are interbedded with the silty mudstones. the sandstones are highly micaceous and may contain abundant carbonaceous detritus. they are massive or show small-scale wave ripple cross-lamination which passes laterally into hummocky cross-stratification. the sandstones are bioturbated; skolithos isp. and taenidium serpentinum are common. body fossils are common in the sandstones but not in the silty mudstones (callomon & birkelund 1980; fürsich 1984). this may reflect the nature of the exposures, but may also reflect primary differences in abundance. interpretation the muds were deposited from suspension fall-out in a marine offshore environment around and below storm wave base, as testified by the fine grain size and the parallel lamination. the sharp-based sandstone beds are interpreted as having been deposited during single storm events whereas the thicker units represent amalgamated deposits formed during successive storms. the well-sorted fine-grained sediment and the small-scale wave ripple cross-lamination associated with hummocky cross-stratified levels suggest deposition on a shoreface between fair-weather and storm wave base (duke 1985). comparison with recent shallow marine environments suggests water depths of 15–30 m (harms et al. 1975, 1982; dott & bourgeois 1982; brenchley et al. 1986, 1993). the sands were probably transported in suspension by storm-induced currents into the otherwise mud-dominated offshore environment. the fine-grained development of the offshore association at visdal in the upper c. tenuiserratum and a. glosense chronozones is also associated with glauconitic horizons, which represent periods of low sedimentation rates. on modern shelves, glauconite is characteristic of sediment-starved offshore environments seawards of the 30 m isobath (blatt et al. 1980; swift & parsons 1995). facies successions two facies successions representing a basinal (eastern) and proximal (western) position, respectively, are described in order to illustrate changes in the depositional environments through time. recognition of the facies successions is based on vertical sections; they are shown in the geological cross-section in figure 18. the eastern basinal area the succession in the eastern, most basinwards position, probably comprises the oldest sediments outcropping in milne land (fig. 4). from below, the succession consists of a poorly exposed unit of the alluvial association overlain by a thick succession of stacked shoreface sandstones. in the better-exposed upper part of the succession three stacked clinoform units cb1–cb3 occur (fig. 18). each of these units represents a seawards shift in facies suggesting that they formed in response to relative falls in sea level (see previous discussion under the prograding wedge association). the stacking of such sand bodies may be controlled by the fixed position of the basin margin or by a change in shoreface gradient due to underlying faults (jerzykiewicz & wojewoda 1986; trincardi & field 1991). the geometries of the clinoform units cb1–cb3 cannot be ascertained due to limited exposure, but it is possible that they represent laterally extensive sheet-like bodies as described from the volgian raukelv formation of jameson land (surlyk & noe-nygaard 1991) and from the calcarenite di gravina formation of italy (pomar & tropeano 2001). the sandstone bodies are truncated by transgressive surfaces of erosion overlain by a ravinement bed (facies 7) containing worn belemnites and vertical burrows of diplocraterion habichi. in cb3, the pebbly sandstone lag is directly overlain by offshore marine mudstones indicating a marine flooding and a marked landwards shift in facies. offshore mud deposition continued throughout the late callovian – middle oxfordian at this locality indicating that the sand-dominated depositional 918 919 10 m 1 km 3 5 12 w es t ea st 2 1 m ud st on e li th ol og y st ru ct ur e tr ac e an d bo dy fo ss ils k ey s ur fa ce s sa nd st on e c on gl om er at e c ry st al lin e ba se m en t pe bb le la g se qu en ce b ou nd ar y (s b) fa ci es a ss oc ia tio ns a llu vi al r oc ky s ho re sh or ef ac e m as si ve lo w o rd er h ig h or de r h ig h or de r h ig h or de r lo w o rd er h ig h or de r lo w o rd er ? ?? a m m on ite be le m ni te sk ol ith os is p. d ip lo cr at er io n ha bi ch i pl an ol ite s is p. pa ra lle lla m in at ed la rg esc al e cr os sbe dd ed tr ou gh cr os sbe dd ed pr og ra di ng w ed ge o ffs ho re c lin of or m s c ov er ed d ir ec tio n of cl in of or m pr og ra da tio n d at um c. c or da tu m c hr on oz on e c ry st al lin e ba se m en t m ar in e flo od in g su rf ac e (f s) c b 4 m ud sa nd m ud sa nd m ud sa nd m ud sa nd m ud sa nd c b 1 c b 2 c b 3 fi g. 1 8. c ro ss -s ec tio n t h ro u gh t h e c h ar co t b u gt a n d l o w er k ap l es lie f o rm at io n s sh o w in g fa ci es a ss o ci at io n s an d s eq u en ce s tr at ig ra p h ic k ey s u rf ac es . t h e cr o ss -s ec tio n i s b as ed o n se le ct ed k ey s ec tio n s (f o r lo ca tio n s, s ee f ig . 1) . system remained in a more landwards position, probably due to a major rise in relative sea level. it is also possible that a barrier formed by a step in the crystalline basement surface was transgressed in the late callovian resulting in a more permanent landwards translation of the sandstone-dominated depositional system, even during a relatively small rise in relative sea level. the western basin margin at visdal, coarse-grained deposits formed by reworking of alluvial and shoreface deposits during an overall middle jurassic transgression dominate the basal part of the charcot bugt formation (fig. 18). the most characteristic deposit is the coarse-grained rocky shore association which rests either directly on the subaerial unconformity formed by the crystalline basement surface or on a submarine ravinement surface developed on top of alluvial sediments. the coarse-grained deposits represent stacked lags formed by wave ravinement during several sea-level cycles and are time transgressive (see kidwell 1989). the marine conglomerates are truncated by numerous erosional surfaces, some of which may have formed in response to relative sea-level changes of regional significance. the coarse-grained amalgamated nature of the succession and evidence of local erosion by rip channels and shoaling waves, however, preclude distinction between local and regional surfaces. the nature of the exposures also precludes direct tracing of the erosional surfaces into the shoreface deposits. a few kilometres seawards of the rocky shore, the deposits are dominated by stacked coarsening-upwards progradational shoreface units forming parasequences or simple sequences (fig. 18). within the shoreface units, the transgressive part is subordinate (arnott 1995) and is only represented by a thin lag conglomerate, which may be overlain by a mudstone unit less than one metre thick. the bulk of the units represent progradational shoreface deposits formed in response to increasing sediment supply or were controlled by changes in relative sea level. the upper part of the charcot bugt formation at visdal shows a marked change from the stacked highstand shoreface units to clinoform unit cb4. it reaches a thickness of 50 m and downlaps onto transgressive mudstones overlying large-scale cross-bedded shoreface deposits (fig. 18). in the most proximal western areas, the lower boundary cuts deeply into the underlying deposits. two stacked fining-upwards units that consist exclusively of tidally-influenced cross-sets (facies 10) dominate the overlying deposits and the incision may mark the position of a tidal channel. these deposits are downlapped by high-angle, tangential clinoforms of a progradational clinoform set. cb4 has a lobate plan geometry and wedges out over a few kilometres towards the south-east and east. the top of clinoform unit cb4 is strongly truncated in the proximal western part where it is overlain by a coarse-grained pebble lag of fluvial affinity. towards the east, the lag fines and grades into a lag deposit of flat, rounded quartzitic pebbles and worn belemnites that is interpreted to have formed by wave winnowing during marine transgression. cb4 is erosionally overlain by coarse-grained shoreface sandstones (mudderbugt member) in the southern part of the visdal valley. following progradation of cb4, the depositional system was drowned and succeeded by offshore marine muds indicating that the sand-dominated depositional system shifted further westwards during the middle oxfordian. the mudstones are characterised by small chondrites isp. burrows and glauconitic horizons, suggesting a sediment-starved, oxygen-restricted environment. biostratigraphic data show that the offshore mudstones of the lower kosmocerasdal member in the eastern outcrop area are contemporaneous with the coarse-grained marine deposits of the charcot bugt formation to the west. the variation in grain size and content of organic material reflect changes in energy regime, sedimentation rates and probably water depths. the coarsening-upwards offshore successions correlate with progradational units in the charcot bugt formation and shallow-water sandstones encountered in the easternmost outcrops correspond to the final progradation of the charcot bugt formation during the c. tenuiserratum chron (fig. 4). in contrast, the most fine-grained levels were deposited during periods of increasing water depth. condensation reflected by high diversity and density of dinocysts and high toc values are seen in mudstones of the q. lamberti ?, upper q. mariae – lower c. cordatum and upper c. tenuiserratum – a. glosense chronozones. bioturbation during these periods was dominated by chondrites isp. sequence stratigraphic model a hierarchy of depositional sequences is recognised in the middle–upper jurassic succession of milne land and is related to two, or possibly three, orders of relative sea-level change. the low order cycle has a duration of c. 30 ma (bathonian–volgian) and corresponds 920 in duration to the second order cycles of vail et al. (1977). the higher order cycles of the charcot bugt formation may correspond to third order cycles, although a genetic background for cycle orders has never been demonstrated (see discussion by miall 1997). the highest order cycles that can be differentiated are related to the internal upbuilding of the clinoform units, although it is difficult to separate sea-level from autocyclic control at this level. low order cycle the jurassic – lowermost cretaceous succession of east greenland has been interpreted within a low-resolution sequence stratigraphic framework (surlyk 1990, 1991). the cycle is bounded below by a regional sequence boundary, which in milne land coincides with the onlap unconformity between the crystalline basement and the middle jurassic sediments. no low order lowstand deposits are preserved along the western basin margin, which was probably bypassed by sediment during most of the early jurassic. a rise in relative sea level is reflected by the bathonian – middle oxfordian westwards onlap onto the crystalline basement and by the marked backstepping of the depositional systems. the charcot bugt formation and the lower part of the kap leslie formation are thus interpreted to represent the transgressive deposits of a low order sea-level cycle. the maximum flooding zone is represented by lower kimmeridgian laminated organic-rich shales of the gråkløft member, kap leslie formation (birkelund et al. 1984). the sequence is bounded at the top by a major unconformity that is recognised within the sandstone-dominated hartz fjeld formation (birkelund et al. 1984; surlyk et al. 1993). the low order depositional cycle probably reflects increasing rates of subsidence from the onset of rifting in the late bajocian to rift climax in the middle volgian followed by decreasing subsidence rates accompanying waning of rifting in the latest volgian – earliest cretaceous. the regional sea-level cycle correlates with the onset, increase and climax of rifting as recognised throughout the northern north sea – north atlantic region (ziegler 1988). high order cycles the high order sequence stratigraphic interpretation is based on facies successions and the identification of stratigraphic surfaces across which occur major shifts in facies. the correlation and interpretation of high order cycles are supported by biostratigraphic data, but due to their low resolution we are not able to document hiati across the identified sequence boundaries. a subdivision of the entire succession into a definite number of high order cycles is hampered by exposure quality and lack of correlation between the eastern area (localities 1, 2) and the western area (localities 3–19). in the following, however, we discuss sedimentary successions that we believe formed as a result of high order relative sea-level changes. the clinoform units cb1–cb4 unconformably overlie marine shoreface sandstones or offshore mudstones and represent periods of rapid progradation introducing relatively coarse-grained sandstones into more basinal environments. the clinoform units are interpreted as having been formed during falling sea level and possibly early lowstand and define high order sequences. the sandstones may thus be placed in the falling stage systems tract and are capped by the sequence boundary (hunt & tucker 1993, 1995; plint & nummedal 2000). the stacking of cb1–cb3 in a relatively basinal position and the relatively steep nature of the clinoforms suggests that the clinoform units were deposited in a shelf-edge position and thus mark the maximum progradation during a fall in relative sea level. the clinoform units are truncated by a transgressive surface of erosion concealing a sequence boundary, and are overlain either by a shoreface unit (cb1, cb2) or directly by offshore mudstones (cb3, cb4). following progradation of cb3, a marked backstepping of the entire depositional system translated the shoreline far to the west of the present-day outcrop area and deposition in milne land was characterised by silty mudstones of the kosmocerasdal member. during a subsequent sea-level fall, the clinoform unit cb4 prograded, but did not reach the former position of the shelf-edge. in the proximal areas, the presence of a distinct pebble lag suggests that a fluvial system was developed during the final phase of progradation. the upper part of cb4, however, was reworked during the following transgression and the top of cb4 represents a coalesced sequence boundary and transgressive surface of erosion. the transgressive shoreface deposits of the mudderbugt member were formed by reworking of the prograding wedge sands in the southern part of the visdal valley. the transgressive surface of erosion forming the upper boundary of the charcot bugt formation is overlain by offshore glauconitic marine mudstones of the kap leslie formation which are interpreted to form the transgressive systems tract of the following sequence. 921 the boundary between the charcot bugt and kap leslie formations appears, at first sight, to represent a single transgressive surface of erosion. detailed sedimentological and biostratigraphical data show, however, that it represents a system of shingled sub-horizontal transgressive surfaces of erosion (figs 4, 18). the transgressive surfaces apparently merge basinwards, as suggested by surlyk (1991) for the boundary between the contemporaneous pelion and fossilbjerget formations in jameson land. the understanding of the stratal geometry is crucial for the sequence stratigraphic interpretations and for the correct correlation of genetically linked depositional systems (posamentier et al. 1992). marine transgressive surfaces of erosion are the most regionally extensive and continuous surfaces in the charcot bugt formation and some may be traced throughout the entire outcrop. this may be due to their high preservation potential and the fact that marine transgressive erosion commonly removes evidence of subaerial and marine surfaces formed during times of fall and lowstand of sea level leaving only a thin lag deposit (plint 1988; posamentier et al. 1992). wave ravinement is able to remove a significant amount (up to 20 m) of the previously deposited succession (demarest & kraft 1987). the sequence stratigraphic model for the charcot bugt – kap leslie formations suggests a marked spatial shift in depocentres of the deposits of the different systems tracts, although each of the facies associations contain facies deposited during both falling and rising relative sea level (fig. 18). the basinwards eastern area is dominated by stacked falling stage/lowstand sandstone bodies erosionally overlain by thin transgressive lags. highstand deposits occur, but are thin. in the proximal western areas near the structurally controlled basin margin, transgressive deposits onlap the crystalline basement. they are remarkably coarse-grained and thickly developed due to the steep gradient rocky shore. alluvial sediments are preserved in topographic lows. these sediments overlie the sequence boundary and are truncated by the transgressive surface of erosion; they thus belong to the lowstand systems tract. off the rocky shore, the succession is dominated by stacked parasequences representing progradational shoreface units referred to the highstand systems tract (fig. 18). contemporaneous successions the charcot bugt formation forms the basin margin correlative of the pelion, fossilbjerget and olympen formations of jameson land. the detailed middle jurassic boreal ammonite zonation allows direct correlation of the successions and the stacking of depositional units can be shown to be broadly similar (fig. 19; engkilde & surlyk 2003, this volume; larsen & surlyk 2003, this volume). progressive basin margin onlap to the west and north in the basin occurred during deposition of the upper bajocian – lower bathonian sequences p2 and p3, accompanied by a northwards shift of the depocentre (p after pelion formation, see engkilde & surlyk 2003, this volume). sequence p3 was deposited during the early bathonian a. arcticus chron and can be correlated with shallow marine sandstones of the basal charcot bugt formation (fig. 19). engkilde & surlyk (2003, this volume) suggested that sequence p3 correlates with sandstones forming a thick transgressive systems tracts on the islands of traill ø and geographical society ø (fig. 1). the backstepping of high order sequences and the formation of thick transgressive deposits at the basin margin are thus characteristic features of the boreal upper bajocian – lower bathonian jameson land basin (engkilde & surlyk 2003, this volume). in milne land, the middle bathonian a. ishmae chronozone is represented by shoreface sandstones in the lower part of the charcot bugt formation and can be correlated with sequence p4 in jameson land. deposition of sequences p4 and p5 of the vardekløft group was characterised by backstepping in response to continued sea-level rise. the transgressive systems tract of sequence p5 is characterised by ammonites of tethyan affinity suggesting that the transgression may be significant outside east greenland, opening for oceanic circulation to the south (callomon 1993; engkilde & surlyk 2003, this volume; alsen & surlyk in press). in milne land, the middle bathonian a. cranocephaloide chronozone is represented by a thick rocky shoreline succession, which can be correlated with highstand deposits of sequence p5. backstepping continued through sequences p6–p8, but was interrupted by a short progradational event reflected in the deposition of the sandy parnas member (highstand deposits of p6) in northern jameson land (heinberg & birkelund 1984; engkilde & surlyk 2003, this volume). the progradation took place near the boundary between the lower callovian c. nordenskjoeldi and p. koenigi chronozones. this may correlate with the marked progradation of cb3 in milne land (fig. 21). the maximum transgression recorded in the vardekløft group is represented by highly condensed deposits of sequence p8 and the overlying condensed mudstone succession of the fossilbjerget formation, which formed 922 during the callovian s. calloviense, k. jason and e. coronatum chrons (surlyk et al. 1973; callomon 1993; engkilde & surlyk 2003, this volume). this transgressive interval correlates with the backstepping of the charcot bugt formation in milne land and deposition of the thick mudstone succession of the lower kosmocerasdal member (fig. 19). a renewed progradational phase occurred in jameson land in the p. athleta chron and is represented by deep-water turbidites of the athene member of the lower olympen formation (larsen & surlyk 2003, this volume). in milne land, the succession is poorly documented but may tentatively be correlated with a thick unfossiliferous sandstone succession exposed at visdal and aldinger elv. the sandstones are overlain by carbonaceous mudstones representing a major drowning of the coarse-grained depositional system. the mudstone succession in milne land is rich in dinoflagellates and is correlated with the c. cordatum chronozone in jameson land. it is represented by a thick mudstone succession in the hades member of the olympen formation (fig. 19). the thick clinoform unit cb4 represents the final progradation of the charcot bugt formation in milne land. it prograded during the late c. cordatum and c. densiplicatum chrons and correlates with coarse-grained deltaic deposits of the upper olympen formation (fig. 19, zeus member; larsen & surlyk 2003, this volume). the close correlation of the successions in milne land and jameson land suggests that they were deposited in a major basin covering the entire area. the facies differences probably reflect the position relative to the basin margins and to the main sediment entry points. the marked change from the bajocian–callovian ramp setting into the oxfordian–kimmeridgian shelfslope setting illustrated by the pelion, fossilbjerget and olympen formations (engkilde & surlyk 2003, this volume; larsen & surlyk 2003, this volume), is thus not significant in the milne land succession where a shallow shelf environment existed throughout the middle and early late jurassic. stratigraphic correlation on formation scale between lithostratigraphic units in east greenland and the north 923 a. rosenkrantzi a. regulare a. serratum a. glosense c. tenuiserratum c. densiplicatum c. cordatum q. mariae q. lamberti p. athleta e. coronatum k. jason s. calloviense p. koenigi c. nordenskjoeldi c. apertum c. calyx c. variabile a. cranocephaloide a. ishmae a. greenlandicus a. arcticus c. pompeckji c. indistinctus c. borealis boreal chronozones chronostratigraphy dinoflagell. cyst assembl. milne land ba jo ci an ba th on ia n c al lo vi an o xf or di an m m m u u u u l l l l m id dl e ju ra ss ic u pp er ju ra ss ic sequence stratigraphy 140 km 200 km milne land jameson land w e s n ass. 8 ass. 7 ass. 6 ass. 5 ass. 4 ass. 1 ass. 3 ass. 2 hareelv fm hades mbolympen fm fossilbjerget fm no exposure zeus mb athene mb p8 p7 p6 p4 p3 p2 p1 p5 pelion fm lower jurassic fine-grained sandstone coarse-grained sandstone mudstone crystalline basement no exposure ? ? cb1 cb2 cb3 cb4 aldinger elv mb kosmocerasdal mb kap leslie fm bays elv mb charcot bugt fm fig. 19. sequence stratigraphic model for the milne land – jameson land successions based on larsen (1995), engkilde & surlyk (2003, this volume), larsen & surlyk (2003, this volume), and this study. atlantic and the northern north sea has been made by numerous authors (e.g. birkelund 1975; larsen 1987; doré 1991; partington et al. 1993; dam & surlyk 1995, 1998). surlyk et al. (1993) and engkilde & surlyk (2003, this volume) demonstrated the close similarities between sequence stacking patterns of the vardekløft formation in east greenland and correlative rocks in the north atlantic region. the coarse-grained falling stage/lowstand sandstone bodies of the charcot bugt formation thus represent an excellent reservoir analogue and may form the basis for developing a new middle jurassic stratigraphic play. conclusions new biostratigraphic data are presented for the middle and upper jurassic succession of milne land based on ammonite collections and palynology. correlation of the coarse-grained sandstones of the charcot bugt with the fine-grained offshore kap leslie formation suggests a genetic depositional relationship with a proximal to distal decrease in grain size. a widespread transgression in east greenland took place following a major uplift period in late early jurassic – earliest middle jurassic times and the coarse-grained amalgamated deposits of the basal charcot bugt formation were deposited in front of a steep rocky shore. estimates of the middle jurassic relative sea-level rise are based on contour maps of the onlap surface between crystalline basement and shallow marine sediments. an overall sea-level rise of more than 300 m is documented, which is independent of sediment supply except for the loading effects. the most characteristic facies is represented by clinoform units up to 52 m thick with compound clinoforms dipping up to 24°. the clinoform units consist of coarse-grained, locally pebbly sandstones forming a progradational wedge deposited during falling relative sea level. the unusually steep clinoforms are probably a function of the coarse grain size and high progradation rates into a relatively deep-water basinal setting. a sequence stratigraphic interpretation involving two orders of sea-level cycles is presented based on spatial changes in depositional systems (facies associations), the nature of bounding surfaces and biostratigrapic data. this model suggests that lateral shifts in depocentres are related to changes in the relative sea level. thus, falling stage/lowstand prograding wedges occur to the east in a relatively distal position and are encased in distal shoreface and offshore deposits. acknowledgements m.l. gratefully acknowledges the carlsberg foundation for funding of a ph.d. stipendium and field work in east greenland (91-0683/20, 92-0505/20, 93-0735/20). the stratigraphic analysis was supported by the danish energy agency, efp93/0010 and 0017. f.s. acknowledges generous support from the carlsberg foundation, the danish natural science research council and norsk hydro a.s. we thank john h. callomon, university college, london for invaluable identification of the ammonites. the work benefited from numerous discussions and suggestions by michael engkilde and gregers dam. we wish to thank the referees tom dreyer and dale leckie for their constructive comments. references aldinger, h. 1935: geologische beobachtungen im oberen jura des scoresbysundes (ostgrönland). meddelelser om grønland 99, 128 pp. alsen, p. & surlyk, f. in press: maximum middle jurassic transgression in east greenland: evidence from new ammonite finds, bjørnedal, traill ø. in: stemmerik, l. & stouge, s. 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(ed.): seismic stratigraphy – applications to hydrocarbon exploration. american association of petroleum geologists memoir 26, 83–97. ziegler, p.a. 1988: evolution of the arctic – north atlantic and the western tethys. american association of petroleum geologists memoir 43, 198 pp. manuscript received 22 july 1997; revision accepted 7 december 2000. 928 appendix 1: dinoflagellate cyst assemblages assemblage 1 ammonite stratigraphy. a. arcticus chronozone (jameson land fauna 9–10; jameson land faunas defined by callomon 1993) or slightly older. samples. samples 409552, 409553 and 409554 were collected from silty mudstones below the horizon containing ammonites of the a. arcticus chronozone at locality 8 in visdal (fig. 6). description. the assemblage comprises crussolia perireticulata, lithodinia cf. reticulata, solisphaeridium ankyleton, pareodinia ‘birkelundia’, sentusidinium pelionense, dichadogonyaulax sellwoodii, durotrigia daveyi, kallosphaeridium hypornatum, pareodinia halosa and sirmiodinium grossii. the assemblage is of low density and diversity. in jameson land, most of these morphologically characteristic species have their stratigraphic base in the c. pompeckji chronozone or lower (fig. 5). in contrast, sirmiodinium grossii first appears in the lowermost a. arcticus chronozone in jameson land. the presence of crussolia perireticulata, atopodinium sp. and especially a. haromense in this assemblage indicates that these species appear at a lower stratigraphic level than in nearby jameson land. the palynological evidence of the age of assemblage 1 is not very clear, but the presence of s. grossii is suggestive of the earliest a. arcticus chron, in accordance with the ammonite data. chronostratigraphy. a. arcticus chronozone, bathonian. assemblage 2 ammonite stratigraphy. the palynomorph assemblages were obtained from samples collected between horizons containing ammonite faunas referred to the a. arcticus (jameson land fauna 9–10) and a. cranocephaloide/c. variabile chronozones (fig. 5; milne land fauna m1 and younger, jameson land fauna 18–21). samples. sample 409455 (409457 barren) from locality 8 in visdal. description. the common occurrence of gonyaulacysta pectinigera, aldorfia aldorfensis and chytroeisphaeridia chytroeides in the sample indicates an earliest c. variabile chron age, on the basis of correlation with data from jameson land. chronostratigraphy. c. variabile chronozone, bathonian. assemblage 3 ammonite stratigraphy. the palynomorph assemblage occurs above ammonite-bearing beds referred to the a. cranocephaloide? chronozone (milne land fauna m1 and jameson land fauna 18–19). samples. sample 234049 is from a mudstone bed overlying sandstones that yielded an ammonite referred to the a. cranocephaloide? chronozone at kosmocerasdal (fig. 1, locality 1; callomon & birkelund 1980). description. this odd assemblage from just one sample is characterised by a flood of ctenidodinium sp.; this event has not been identified in jameson land. the appearance of evansia granulata supports the age indicated by the ammonite from the underlying beds as it has its first occurrence in the a. cranocephaloide chronozone in jameson land; the underlying a. ishmae chronozone has not, however, been studied in detail. the minimum age of this assemblage is poorly constrained both on the basis of ammonites and dinoflagellate cysts, but kallosphaeridium inornatum and lithodinium cf. reticulata, both of which are present in the assemblage, occur in the uppermost part of the p. koenigi chronozone in jameson land. chronostratigraphy. a. cranocephaloide/c. variabile chronozones, bathonian. assemblage 4 ammonite stratigraphy. sample 255155 was collected from above beds referred to the a. cranocephaloide/c. variable chronozones on the basis of ammonites (jameson land fauna 19–21). samples. samples 255148, 255155 and 409556 are from locality 1 at the base of clinoform unit cb2 in kosmocerasdal and locality 8 in visdal (fig. 1). sample 409556 is from a mudstone horizon above, but close to an ammonite horizon indicating the a. cranocephaloide/c. variabile chronozones. samples 255155 and 255148 are thought to be from the same horizon but are separated laterally and direct correlation is not possible. description. the diversity of the assemblage varies but abundant chyctroeisphaeridia hyalina, evansia granulata, sentusidinium pelionense and lithodinia cf. reticulata characterise the microflora. the appearance of abundant new species suggests a clear separation in time from the underlying assemblage. chytriasphaeridia hyalina, atopodinium polygonalis and evansia granulata become abundant for the first time in the c. calyx chronozone in 929 jameson land. evansia janeae occurs only in the c. calyx chronozone in jameson land in contrast to the occurrence in milne land. lithodinia spongiosa is restricted to the c. calyx and c. apertum chronozones in jameson land. a number of characteristic species appearing in this assemblage, viz. evansia cerebraloides, paragonyaulacysta sp. (cf. calloviense, baylei et al.), kalyptea stegasta, sentusidinium sp. d (fensome 1979), ctenidodinium thulium and meiourogonyaulax cf. planoseptata, are considered stratigraphically significant but cannot yet be correlated precisely with the succession in jameson land. chronostratigraphy. c. calyx/c. apertum chronozones, bathonian. assemblage 5 ammonite stratigraphy. none. samples. three samples in stratigraphic succession from the bottomset of clinoform unit cb3 at kosmocerasdal (locality 1): 409697, 409698 and 409699. description. abundant new species appear in assemblage 5, again indicating a clear time difference to the underlying assemblage. the assemblage is of high density and diverse in contrast to the earlier assemblages. sirmiodinium grossii, g. pectinigera, lithosphaeridium cf. reticulata, atopodinium sp., evansia cerebraloides, ctenidodinium thulium and sentusidinium sp. d (fensome 1979) become abundant and rhynchodiniopsis cladophora, gonyaulacysta cf. helicoidea, gonyaulacysta jurassica, fromea tornatilis, chytroeisphaeridia chytroeides, ellipsoidictyum cinctum and lithodinium jurassica appear for the first time. approximately 50% of the species in assemblage 5 occur in the earlier assemblages but not with this high abundance. only a few of the new species in the assemblage are abundant. local correlation. abundant g. pectinigera has an upper limit in the basal p. koenigi chronozone in jameson land; this limit is recognised within this assemblage. crussolia perireticulata has an apparent top in the c. nordenskjoeldi chronozone in jameson land and this may also be recognised within this assemblage. chytroeisphaeridia chytroeides is almost wholly restricted to this interval, and is similarly present or abundant in the c. nordenskjoeldi chronozone in jameson land. atopodinium polygonalis is only recorded in the c. calyx and c. apertum chronozones in jameson land, whereas paraevansia brachythelis is recorded in the c. apertum and c. nordenskjoeldi chronozones. a suite of species which are uncommon in the assemblage have their stratigraphic tops in the basal p. koenigi chronozone in jameson land. chronostratigraphy. c. nordenskjoeldi – earliest p. koenigi chronozone, lower callovian. assemblage 6 ammonite stratigraphy. the assemblage occurs beneath ammonite-bearing beds referred to the p. athleta chronozone. samples. two samples (255153, 255154) in the basal strata of the kap leslie formation in kosmocerasdal (locality 1). description. the assemblage is characterised by medium diversity and low abundance. chytroeisphaeridia hyalina is the only abundant species accompanied by pareodinia prolongata in one sample. only few stratigraphically new species appear in the assemblage suggesting stratigraphic proximity to the underlying assemblage 5. local correlation. the ammonite and dinoflagellate stratigraphy in the jameson land basin is not well correlated in the interval between the s. calloviense and p. athleta chronozones (fauna m2), but the distribution of dinoflagellate cysts in both chronozones is fairly well known. in jameson land, chytroeisphaeridia hyalina occurs in abundance in the s. calloviense chronozone, coincident with the last continuous occurrence of gonyaulacysta pectinigera. the only abundant dinoflagellate recorded by fensome (1979) from the s. calloviense chronozone of jameson land was valensiella dictydia (15–30% of the assemblage), associated with gonyaulacysta pectinigera and valensiella ovula (5–15% of the assemblage). v. dictydia and v. ovula have been recorded previously from these strata (piasecki 1980; smelror 1988) but were not observed in the present samples. fensome (1979) also described paragonyaulacysta sp. (possibly equivalent to paragonyaulacysta retiphragmata) from this chronozone and his species is recorded here. the assemblage is correlated with the s. calloviense chronozone. chronostratigraphy. s. calloviense chronozone, lower callovian. assemblage 7 ammonite stratigraphy. none. samples. samples (255147, 255149–255151, 409543, 409544, 409550, 409558) from the bottomset of clinoform unit cb4, in the uppermost charcot bugt formation in visdal (localities 8, 9, 11). description. the dinoflagellate assemblage is dominated by gonyaulacysta jurassica, rhynchodiniopsis cladophora, liesbergia 930 scarburghensis, rigaudella aemula and pareodinia stegasta at successively higher stratigraphic levels. local correlation. the presence of wanaea thysanota in association with rigaudella aemula, liesbergia scarburghense and scriniodinium crystallinum corresponds to the interval between ammonite faunas m4 and m5 of milne land (c. cordatum to lowermost c. densiplicatum chronozones) in the fine-grained kap leslie formation. the samples in the same bed towards the east show the top of abundant rigaudella aemula, followed by an acme of pareodinia stegasta. these two events correspond to the level between milne land ammonite faunas 6 and 7, indicative of the upper c. densiplicatum chronozone. chronostratigraphy. c. cordatum to c. densiplicatum chronozones, lower–middle oxfordian. assemblage 8 ammonite stratigraphy. none. samples. samples (409546, 409551, 409559) from the basal part of the kap leslie formation onlapping the charcot bugt formation, from the east towards the west (fig. 1, localities 5, 12). description. dinoflagellate cysts from two samples of the basal mudstones (409546 is effectively barren) reflect the stratigraphic onlap. the common presence of ambonosphaera calloviana and dingodinium sp. indicate a level corresponding to milne land ammonite fauna 10 (upper a. glosense chronozone), and in the succeeding sample, epiplosphaera bireticulata and stephanelytron redcliffense appear and indicate the stratigraphic level above milne land ammonite fauna 10 (a. glosense to a. serratum chronozones). chronostratigraphy. a. glosense to a. serratum chronozones, upper oxfordian. geological survey of denmark and greenland bulletin 1, 437-458 north sea baltic sea 10°e8°e 12°e 14°e 16°e 58°n 57°n 56°n 55°n 6°e4°e sweden the netherlands germany norway denmark uk danish basin fault national boundaries bornholm 50 kmfennoscandian border zone danish basin and fennoscandian border zone 439 the tornquist zone is a fundamental tectonic lineament representing the south-western margin of the baltic shield (fig. 1). the lineament runs north-west from the carpathians across poland, where it is known as the teisseyre–tornquist zone, and into the scandinavian area, where it is known as the sorgenfrei–tornquist zone (eugeno-s working group 1988). it crosses northern denmark in a nw–se direction and extends as far as the viking graben in the north sea (pegrum 1984). the lineament had its origin in precambrian times and faults defining the lineament have been intermittently active until the present day. it is characterised by complex extensional and strike-slip faulting and structural inversion (liboriussen et al. 1987; eugeno-s working group 1988; mogensen 1992a, b, 1994; mogensen & korstgård 1993; christensen & korstgård 1994; mogensen & jensen 1994). the sorgenfrei–tornquist zone within and adjacent to the kattegat area (fig. 1), has been described in several papers, based on field information (bergström et al. 1982; norling & bergström 1987; sivhed 1991), well data (michelsen & nielsen 1991, 1993; jensen & michelsen triassic and jurassic transtension along part of the sorgenfrei–tornquist zone in the danish kattegat tommy egebjerg mogensen and john a. korstgård in the kattegat area, denmark, the sorgenfrei–tornquist zone, an old crustal weakness zone, was repeatedly reactivated during triassic, jurassic and early cretaceous times with dextral transtensional movements along the major boundary faults. these tectonic events were minor compared to the tectonic events of the late carboniferous – early permian and the late cretaceous – early tertiary, although a dynamic structural and stratigraphic analysis indicates that the sorgenfrei–tornquist zone was active compared to the surrounding areas. at the end of the palaeozoic, the area was a peneplain. regional triassic subsidence caused onlap towards the north-east, where the youngest triassic sediments overlie precambrian crystalline basement. during the early triassic, several of the major early permian faults were reactivated, probably with dextral strike-slip along the børglum fault. jurassic – early cretaceous subsidence was restricted primarily to the area between the two main faults in the sorgenfrei–tornquist zone, the grenå–helsingborg fault and the børglum fault. this restriction of basin development indicates a change in the regional stress field at the triassic–jurassic transition. middle jurassic and late jurassic – early cretaceous subsidence followed the early jurassic pattern with local subsidence in the sorgenfrei–tornquist zone, but now even more restricted to within the zone. the subsidence showed a decrease in the middle jurassic, and increased again during late jurassic – early cretaceous times. small faults were generated internally in the sorgenfrei–tornquist zone during the mesozoic with a pattern that indicates a broad transfer of strike-slip/oblique-slip motion from the grenå–helsingborg fault to the børglum fault. keywords: kattegat, denmark, sorgenfrei–tornquist zone, triassic–jurassic, børglum fault, grenå–helsingborg fault, transtension, structural evolution t.e.m., norsk hydro a.s., n-0246 oslo, norway. e-mail: tommy.mogensen.egebjerg@hydro.com j.a.k., geological institute, university of aarhus, c.f. møllers allé 120, dk-8000 århus c, denmark. geological survey of denmark and greenland bulletin 1, 439–458 (2003) © geus, 2003 440 1992), deep reflection – refraction seismic data (eugeno-s working group 1988; lie & husebye 1992), and reflection seismic data (baartman & christensen 1975; bergström 1984; pegrum 1984; liboriussen et al. 1987; ziegler 1987, 1990; aubert 1988; bergström et al. 1990a, b; ro et al. 1990a; vejbæk 1990). the majority of these studies have proposed lateral movements along the sorgenfrei–tornquist zone with right-lateral movements during the palaeozoic and generally left-lateral movements during the mesozoic (bergström et al. 1982; pegrum 1984; liboriussen et al. 1987; norling & bergström 1987; aubert 1988; sivhed 1991). many of the earlier kattegat studies based on reflection seismic data suffered from large line spacing (pegrum 1984; aubert 1988). the average spacing prior to the seismic surveys from the mid-1980s was around 10 km, which severely hampered detailed structural interpretation such as fault correlation and depocentre configuration along the sorgenfrei–tornquist zone. interpretation of closely-spaced 2d reflection seismic data (1 km spacing in the middle of kattegat, fig. 2), released to danish research institutions in the early 1990s, and tz tz 100 km 10˚e6˚e 14˚e 55˚n sweden norway og børglum faultfjerritslev fault 59˚n 57˚n tz 2 denmark jylland skåne fb sg eb 1 kattegat north sea grenå– helsingborg fault fig. 1. southern scandinavia showing the sorgenfrei–tornquist zone and the area of study (outlined areas 1 and 2 – area 2 was investigated in particular detail). eb, egersund basin; fb, farsund basin; og, oslo graben; sg, skagerrak graben; tz, sorgenfrei– tornquist zone. 1 2 6 3 4 40 km denmark a 1 2 6 3 4 1 2 6 3 4 seismic surveys and wells kattegat sweden norway skåne 10˚e 11˚e 12˚e9˚e 57˚n 56˚n 555 thisted-4 thisted-2 fjerritslev-1,2 hyllebjerg-1 farsø-1 aars-1 vedsted-1 haldager-1 flyvbjerg-1 børglum-1 sæby-1 skagen-2 frederikshavn-1 hobro-1 voldum-1 rønde-1 hans-1 lavø-1 anholt-(1–4) felicia-1 j-1 terne-1 gassum-1 fig. 2. seismic surveys and wells used. the geoseismic profiles in figure 3 are numbered 1–6; the positions of these key profiles are also shown on all subsequent maps (fig 4–13). the backstripped section in figure 14 is indicated by the dashed line a. new well data (michelsen & nielsen 1991, 1993) has made it possible to undertake a more detailed analysis of the palaeozoic and cretaceous structural developments along the sorgenfrei–tornquist zone (mogensen 1992a, b, 1994; mogensen & korstgård 1993; mogensen & jensen 1994). the remainder of the mesozoic, from the triassic to the lower cretaceous, with emphasis on the jurassic, is the scope of this study. much of the discussion in the following is focussed on the two main faults of the sorgenfrei–tornquist zone in the kattegat area, the børglum fault and the grenå– helsingborg fault. these faults, crossing the area from skåne, sweden to north jylland, denmark (fig. 1), are considered as two separate strands of the sorgenfrei– tornquist zone. data this study is primarily based on released 2d reflection seismic data, acquired during a period of hydrocarbon exploration in the area in the early 1980s. exploration also included drilling of the first two deep wells, hans-1 and terne-1 in the central part of kattegat, and the sæby-1 well in north jylland, all penetrating mesozoic rocks (fig. 2; michelsen & nielsen 1991). the reflection seismic surveys used in this study vary in quality. there is a progressive increase in quality from the onshore single fold seismic, shot in 1967, to the 60 fold seismic data shot in 1985 (table 1). resolution of the seismic data is higher in the offshore data, but on a few onshore regional lines continuous reflections down to 4 seconds two-way travel time (twt) can be seen. data from all wells in the area (table 2), and information from rocks outcropping in skåne, south-west sweden, have been used in the study. in the central kattegat area and onshore denmark, there is a good tie between well data and the reflection seismic surveys. elsewhere the interpretation can be more speculative, because of the scarcity of high resolution seismic data (fig. 2). the location of six key lines, shown in figure 3, is indicated on all maps. the study has resulted in a series of maps (figs 4–13); the structure maps (in twt) represent top pre-zechstein (approximately equivalent to the base triassic over most of the investigated area), base jurassic and base cretaceous. isochore maps (in twt) have been prepared for the lower and upper triassic successions, the gassum formation, the fjerritslev formation, the middle jurassic, the upper jurassic and the lower cretaceous successions. where resolution of the onshore surveys is good, the maps cover both the kattegat area and the onshore area (fig. 1, areas 1 and 2); where onshore resolution is poor, the maps only cover the kattegat area (fig. 1, area 2). 441 survey year fold filtered/ onshore/ migrated offshore ssl6267 1964–7 1 filtered onshore wgc67a 1967 6 filtered offshore prkl7374a 1973–4 6–12 filtered onshore gsi75b 1975 12 migrated* offshore dnj8183d 1982–3 12 migrated onshore rtd81k 1982 48 migrated offshore dcs81k 1982 48 migrated offshore gy82k 1983 48 migrated offshore geco83ak 1983 48 migrated offshore dn84d 1984 24 migrated onshore dk84k 1984 48 migrated offshore am84k 1984 60 migrated offshore tx84k 1984 48 migrated offshore ao85i 1985 24 migrated onshore tx85k 1985 48 migrated offshore * reprocessed and migrated 1983. table 1. seismic surveys used in the studies table 2. study wells grouped into those penetrating the mesozoic (a) and those terminating in the mesozoic (b) a well encounters/terminates in frederikshavn-1 mesozoic/precambrian gassum-1 mesozoic/zechstein hans-1 mesozoic/upper carboniferous rønde-1 mesozoic/upper silurian sæby-1 mesozoic/rotliegend terne-1 mesozoic/cambrian thisted-4 mesozoic/zechstein b well terminates in aars-1 upper triassic børglum-1 upper triassic – lower jurassic farsø-1 upper triassic fjerritslev-1 lower jurassic fjerritslev-2 upper triassic flyvbjerg-1 upper triassic frederikshavn-2, -3 triassic haldager-1 lower jurassic hobro-1 upper triassic hyllebjerg-1 upper triassic lavø-1 upper triassic skagen-2 triassic thisted-2 lower triassic vedsted-1 upper triassic voldum-1 upper triassic data from nielsen & japsen (1991). s sæby-1 ? ? ? ghf bf dnj-400 sorgenfrei–tornquist zone gassum-1w e ssw nne ghf bf dnj-30/k84-002 6.64 cm sorgenfrei–tornquist zone sw nehans-1 (proj.) ghf bf dcs-68 sorgenfrei–tornquist zone ? sw ne ghf bf (a) bf (b) rtd 81-k10/k75-014 sorgenfrei–tornquist zone anholt (1–4) (proj.) 0 2 4 0 2 4 0 2 4 0 2 4 0 2 4 0 2 4 0 2 4 0 2 4 n 5 km 5 km 5 km 5 km 1 2 3 4 t w t s ec t w t s ec t w t s ec t w t s ec t w t s ec 0 2 4 0 2 4 rtd 81-k11 n sorgenfrei–tornquist zone s ghf 5 bf (a) bf (b) 5 km 443 u. cretaceous u. jurassic – l. cretaceous m. jurassic (haldager fm) l. jurassic (fjerritslev fm) u. triassic – l. jurassic (gassum fm) l. triassic – u. triassic basement grenå–helsingborg fault børglum fault ghf bf zechstein salt rotliegendes syn-rift sequence l. palaeozoic n ghf rtd 81-k12 sorgenfrei–tornquist zone t w t s ec 0 2 4 0 2 4 s 6 bf (a) bf (b) 5 km facing page and above: fig. 3. geoseismic profiles (fig. 2 for locations). note: (1) the late cretaceous anticline with underlying depocentres on profile 2, (2) the extensive mesozoic small-scale faulting, especially on profiles 4 and 6, (3) the increased jurassic subsidence in the sorgenfrei–tornquist zone on profiles 3 and 5, and (4) the thinning of the triassic towards the north-east on profiles 1, 2 and 6. note also fault strands a and b of the børglum fault (bf) on profiles 3, 5 and 6, where the a strand seems to take up the lateral component, whereas the b strand seems to take up any vertical component (see text). note that profile 2 is constructed from two seismic lines roughly at right angles to each other and that the two lines do not intersect (fig. 2). jurassic triassic precambrian basement u. cretaceous quaternary subcrop rocks onshore sweden 20 km ghf bf a b well fault 1500–2000 2000–2500 2500–3000 0–500 500–1000 1000–1500 3000–3500 11°e 12°e 56°n 1 2 3 4 5 6 a top pre-zechstein structure map depth in msec twt hans-1 terne-1 fig. 4. top pre-zechstein twt structure map. note that the jurassic onlaps partly onto the basement, partly onto the triassic in skåne, indicating younging of the sediments towards the north-east. note also the zone of lateral transfer between the hans-1 and terne-1 wells. bf, børglum fault, a and b are strands of the børglum fault; ghf, grenå–helsingborg fault. structural development triassic during the early permian, the kattegat area and the sorgenfrei–tornquist zone were exposed to erosion, and prior to the late permian the area had become a peneplain (figs 3, 4; michelsen & nielsen 1991; mogensen 1994). the top pre-zechstein twt structure map (fig. 4), shows the configuration of this peneplain surface today, formed by the sum of all post-early permian tectonic events. the predominant fault orientation is nw–se, the same as the main trend during the palaeozoic, but several nne–sswto nnw–sse-trending faults are also present, both within and outside the sorgenfrei–tornquist zone. after formation of the peneplain, the area started to tilt towards the south-west, with the two main faults in the kattegat, the grenå–helsingborg fault and the børglum fault (fig. 1), lowering the top pre-zechstein surface stepwise towards the south-west (figs 3, 4). in the south-east, the grenå–helsingborg fault is the dominant fault, whereas the børglum fault takes over in an én echelon fashion towards the north-west, where it becomes the main mesozoic fault. in the area between the hans-1 and terne-1 wells and the two major faults (fig. 4), many smaller faults cut the top pre-zechstein surface. this intense faulting occurs to the north-west of a large bend in the grenå–helsingborg fault, and seems to be linked to this bend. north-west of the terne-1 well, the grenå–helsingborg fault gradually 444 20 km ? ? ? ? ghf bf a b 11˚e 12˚e 57˚n 56˚n 1 2 5 6 3 4 a lower triassic isochore map thickness in msec twt 0–100 100–200 200–300 300–400 lower triassic not present (non-deposition or erosion) 400–500 500–700 700–800 fault inferred fault well terne-1 lavø-1 fig. 5. twt isochore map of the lower triassic. note the subsidence pattern along the børglum fault with maxima and minima, indicating strike-slip motion along this fault with push-up and pulldown. note also that there is no general differential subsidence within the sorgenfrei–tornquist zone, the zone is merely the eastward limit of the regional triassic basin to the south-west. bf, børglum fault, a and b are strands of the børglum fault; ghf, grenå–helsingborg fault. 445 dies out, and terminates close to profile 2 (fig. 3). the børglum fault also fades out, much further to the northwest, but no clear termination can be seen (fig. 4). some branches of the børglum fault are relatively straight, and can be followed continuously across the kattegat for 80–100 km. the top pre-zechstein twt structure map (fig. 4) also incorporates a quaternary subcrop map of westernmost skåne, sweden. note that the jurassic onlaps partly on precambrian crystalline basement and partly on the triassic, indicating younging of sediments towards the north-east. this is clearly seen on the interpreted seismic sections which show seismic onlap towards the north-east (fig. 3), where the youngest triassic sediments overlie precambrian basement. the twt isochore maps of the lower triassic (fig. 5) and the middle–upper triassic (fig. 6) illustrate the structural history during the triassic. both maps and the geoseismic profiles (fig. 3) show the general subsidence towards the south-west, where one of the main regional triassic depocentres of the northwest european basin is located (bertelsen 1980; vejbæk 1990). the triassic twt isochore maps indicate that only minor differential subsidence took place along the sorgenfrei– tornquist zone during the triassic. the zone was merely the north-eastward limit of the large triassic basin. during the early triassic, some of this minor differential subsidence internally in the sorgenfrei–tornquist zone took place along bends of the børglum fault (fig. 5), whereas other bends only show limited deposition or ? ? ? ? ghf bf a b 20 km 11˚e 12˚e 57˚n 56˚n 1 2 5 6 3 4 a middle–upper triassic isochore map thickness in msec twt 0–100 100–200 300–400 200–300 middle and upper triassic not present (non-deposition or erosion) 400–500 500–700 700–900 fault well terne-1 lavø-1 fig. 6. twt isochore map of the middle– upper triassic. note the diminishing of maxima and minima along the børglum fault, indicating less activity along this fault compared to the early triassic. bf, børglum fault, a and b are strands of the børglum fault; ghf, grenå–helsingborg fault. erosion. this differential subsidence along the børglum fault abated during the middle and late triassic (fig. 6). west of the terne-1 well, the grenå–helsingborg fault seems to have had only limited effect on deposition (figs 5, 6), whereas the børglum fault had a more pronounced impact on the depositional pattern further to the north-west (figs 5, 6). towards the end of the triassic, the depositional pattern changed and differential subsidence within the sorgenfrei–tornquist zone started to overshadow regional subsidence with deposition of the upper triassic – lower jurassic gassum formation (fig. 7). the børglum fault was still the most active fault in the kattegat area, with a variable subsidence pattern along strike (fig. 7), although the grenå–helsingborg fault seems to have had some influence on the depositional pattern in the western part of the kattegat, close to profile 2 (fig. 7). the gassum formation has a diachronous upper boundary in the danish area, younging towards the basin margin (michelsen & nielsen 1991) which cannot be seen on the seismic lines due to limited thicknesses outside the central kattegat area. jurassic the differential subsidence within the sorgenfrei–tornquist zone that started in the late triassic was enhanced during the early jurassic, as indicated by the twt isochore map of the lower jurassic fjerritslev formation (fig. 8). the thick lower jurassic succession has been deeply eroded along the børglum fault, primarily along fault strand b (figs 3 (profile 3), 8), due to subsequent late cretaceous – early tertiary inversion along this fault 446 ghf bfa b 40 km 10˚e 11˚e 12˚e 9˚e 57˚n 56˚n 200–300 300–400 0–100 100–200 eroded sequence upper triassic – lower jurassic (gassum fm) isochore map thickness in msec twt wellnormal fault limit of gassum fm1 2 5 6 3 4 a terne-1 fig. 7. twt isochore map of the gassum formation. note the onset of differential subsidence in the sorgenfrei–tornquist zone, with the børglum fault being the most active. the gassum formation is the youngest unit that is not affected by erosion caused by the late cretaceous – early tertiary inversion and neogene uplift. bf, børglum fault, a and b are strands of the børglum fault; ghf, grenå–helsingborg fault. strand. the lower jurassic succession thickens towards the eroded area, making the børglum fault the main fault at this time in the central kattegat area. this contrasts with the triassic faulting, when fault strand a was the most active (figs 3 (profile 3), 5, 6). the grenå– helsingborg fault seems to have had the same influence on the early jurassic depositional pattern in the western part of kattegat close to profile 2, as during deposition of the gassum formation (figs 7, 8). the differential subsidence within the sorgenfrei– tornquist zone continued during deposition of the middle jurassic haldager sand formation (fig. 9), although subsidence seems to have decreased compared to the early jurassic. only minor fault activity took place, and the succession is thin with a rather uniform thickness, although a minor depocentre was formed around the terne-1 well. the haldager sand formation may also have increased in thickness towards the børglum fault, as did the gassum formation and presumably the fjerritslev formation, but has later been removed by erosion. increased deposition compared to the middle jurassic is seen on the twt isochore map of the upper jurassic frederikshavn, børglum and flyvbjerg formations (fig. 10). the differential subsidence within the sorgenfrei–tornquist zone can also be seen on the isochore map, with a thickening of these successions towards the børglum fault, although most of these deposits were later removed by erosion following late cretaceous – early tertiary inversion. late jurassic subsidence patterns continued into the early cretaceous (fig. 11), and small local depocentres developed within the sorgenfrei–tornquist zone, coinciding with the lower jurassic depocentres (fig. 8) and with inversion highs formed during the late cretaceous 447 ghf bfa b 10˚e 11˚e 12˚e 9˚e 57˚n 56˚n 40 km 0–100 200–300 100–200 300–400 removed by erosion thickness in msec twt 400–500 500–600 600–700 eroded sequence lower jurassic (fjerritslev fm) isochore map wellnormal fault salt diapir 1 2 5 6 3 4 a fig. 8. twt isochore map of the fjerritslev formation. note differential subsidence all along the sorgenfrei–tornquist zone; the unit is eroded along the børglum fault. note also the depocentre at profile 2 and line a, which might have been caused by dextral transtensional sagging at the termination of the grenå–helsingborg fault (see text). bf, børglum fault, a and b are strands of the børglum fault; ghf, grenå–helsingborg fault. 448 ? ? g h f bf a b 20 k m 11 ˚e 12 ˚e 57 ˚n 56 ˚n 1 2 5 6 3 4 a m id dl e ju ra ss ic (h al da ge r sa nd f m ) is oc ho re m ap t hi ck ne ss in m se c t w t m id dl e ju ra ss ic no t pr es en t (n on de po si tio n or e ro si on ) m id dl e ju ra ss ic p ar tly re m ov ed b y er os io n fa ul t w el l 0– 10 0 10 0– 20 0 t er ne -1 fi g. 9 . t w t i so ch o re m ap o f th e h al d ag er s an d f o rm at io n . n o te t h e lo ca l d ep o ce n tr e ar o u n d t er n e1 an d d ep o si tio n r es tr ic te d t o w ith in t h e so rg en fr ei –t o rn q u is t z o n e. b f , b ø rg lu m f au lt, a an d b ar e st ra n d s o f t h e b ø rg lu m f au lt; g h f , g re n å– h el si n gb o rg f au lt. ? ? g h f bf a b 20 k m 11 ˚e 12 ˚e 57 ˚n 56 ˚n 1 2 5 6 3 4 a u pp er ju ra ss ic (f re de ri ks ha vn , b ør gl um a nd fl yv bj er g fm s) is oc ho re m ap t hi ck ne ss in m se c t w t u pp er ju ra ss ic a bs en t (n on de po si tio n or e ro si on ) u pp er ju ra ss ic p ar tly re m ov ed b y er os io n fa ul t w el l 0– 10 0 10 0– 20 0 20 0– 30 0 30 0– 40 0 t er ne -1 fi g. 1 0. t w t i so ch o re m ap o f th e u p p er j u ra ss ic s u cc es si o n . n o te t h e in cr ea se d s u b si de nc e re la tiv e to th e m id dl e ju ra ss ic , a nd th e in cr ea si ng e ff ec t o f e ro si on . t he d ep os iti on al p at te rn i n d ic at es a t h ic ke n in g to w ar d s th e st ro n gl y in ve rt ed b ø rg lu m f au lt, w h er e th e u p p er j u ra ss ic i s re m o ve d b y er o si o n . b f , b ø rg lu m f au lt, a a n d b ar e st ra n d s o f th e b ø rg lu m f au lt; g h f , g re n å– h el si n gb o rg f au lt. 449 – early tertiary inversion (mogensen & jensen 1994). one of these structures can be seen on profile 2 (fig. 3) and on line a (see also fig. 14). on the top triassic twt structure map (fig. 12), between the terne-1 and hans-1 wells, a large number of faults occur, compared to the top pre-zechstein structure map (fig. 4). these are small-scale faults that are restricted to the mesozoic succession (fig. 3, profile 4). on the base cretaceous twt structure map (fig. 13), this faulting is not indicated, due to the deep erosion of the mesozoic sequence, but it is clearly seen that the børglum fault continues much further to the north than the grenå–helsingborg fault. this reflects a transfer of lateral movement from the grenå–helsingborg fault to the børglum fault from south-east to north-west and is probably related mainly to late cretaceous – early tertiary dextral transpression. mesozoic dextral transtensional structural development triassic rather than being primarily a zone of differential subsidence during the triassic, as suggested by michelsen & nielsen (1991, 1993), we consider the sorgenfrei–tornquist zone in the kattegat area to represent a staircase stepping down from the north-eastern platform to the deep basin in the south-west, delineating the triassic northwest european basin (figs 5, 6). differential subsidence seen internally in the sorgenfrei–tornquist zone to the north-west, along the fjerritslev fault (vejbæk 1990), might be due to salt withdrawal from a zechstein salt basin along the fjerritslev fault (christensen & korstgård 1994) triggered by minor triassic reactivations of this old fault (norling & bergström 1987). this would ? ghf bf b 40 km 10˚e 11˚e 12˚e 9˚e 57˚n 56˚n 1 2 5 6 3 4 a 400–800 800–12000–400 lower cretaceous partly eroded lower cretaceous isochore map thickness in msec twt wellnormal fault limit of lower cretaceous salt diapir fig. 11. twt isochore map of the lower cretaceous succession. note the similarity with the lower jurassic depositional pattern (fig. 8), as well as the increased effect of later erosion. bf, børglum fault, b is a strand of the børglum fault; ghf, grenå–helsingborg fault. 450 imply that no major differential subsidence related to strike-slip faulting took place along the sorgenfrei– tornquist zone during the triassic, as suggested by pegrum (1984), ziegler (1987, 1990) and the eugeno-s working group (1988). the intense rifting and fault-controlled subsidence of the nne–ssw-oriented horn graben (vejbæk 1990; clausen & korstgård 1993, 1994), indicates a stress field in which the least principal stress axis had a wnw–eseorientation. several normal faults in the kattegat area, outside the sorgenfrei–tornquist zone, are in accordance with such a stress field e.g. north of the lavø-1 well, (figs 5, 6), and the nne–ssw-trending svedala fault in skåne, sweden (norling & bergström 1987). if such a wnw–ese-oriented extensional regime existed in the triassic in the kattegat area, the nw–se-trending grenå– helsingborg, børglum and fjerritslev faults would experience right-lateral, probably transtensional strike-slip, motion along their fault planes, which is in agreement with vejbæk (1990). right-lateral transtensional reactivation of the fjerritslev fault, would also favour halokinetic movements and differential subsidence (koyi & petersen 1993) resembling rifting along this fault. the depositional pattern with thin and thick successions related to bends in the børglum fault indicates lateral motion along the faults, with push-up at restraining bends and pull-down at releasing bends (aydin & nur 1985; christie-blick & biddle 1985; harding ghf bf a b 20 km 11˚e 12˚e 57˚n 56˚n 1 2 5 6 3 4 a top triassic (base gassum fm) structure map depth in msec twt 0–300 300–600 600–900 900–1200 1200–1500 1500–1800 normal fault reverse fault well terne-1 hans-1 fig. 12. top triassic twt structure map. note the large number of faults between terne-1 and hans-1, transferring the strike-slip/oblique-slip motion. note also the horsetail splays at several locations. bf, børglum fault, a and b are strands of the børglum fault; ghf, grenå– helsingborg fault. 451 et al. 1985). this restraining–releasing bend pattern was especially active during the early triassic (fig. 5, see also mogensen 1994), and primarily along strand a of the børglum fault (profile 3, fig. 3). strand a resembles the experiments on reactivation of basement faults under conditions of oblique-slip carried out by richard (1991). this movement pattern decreased in activity during the rest of the triassic (fig. 6), whereas the limited deposition to the north-east of the terne-1 well persisted throughout the triassic (figs 5–7). the area of limited sedimentation coincides with the heavily faulted area between the terne-1 and hans-1 wells (figs 4, 12). this fault pattern indicates a broad transfer of strikeslip or oblique-slip motion from the grenå–helsingborg fault to the børglum fault, corresponding to the gradual termination of the grenå–helsingborg fault towards the north-west. during the triassic, deposition gradually overstepped the margins of the regional basin, with progressive onlap towards the north-east (figs 3, 5, 6; norling & bergström 1987; eugeno-s working group 1988). while the triassic was characterised by regional subsidence (figs 5, 6; vejbæk 1990), the jurassic – early cretaceous was dominated by differential subsidence in the area between the two main faults in the sorgenfrei–tornquist zone, the grenå–helsingborg fault and the børglum fault. the change in this local basin development indicates a shift in the regional stress field (see also norling & bergström 1987) probably related to the widespread late triassic – early jurassic rift phase in northwest europe (ziegler 1990). the change in the stress field occurred during deposition of the gassum formation (fig. 7), and the lower jurassic fjerritslev formation (fig. 8), both relatively widespread formations that increase in thickness towards the sorgenfrei–tornquist zone. the sub-regional subsidence around the sorgenfrei– tornquist zone abated during deposition of the shallow 12• 10˚e 11˚e 12˚e 9˚e 57˚n 56˚n 40 km ? ghf bf a b bf 400–800 800–12000–400 1200–1600 base cretaceous structure map depth in msec twt wellnormal fault reverse fault limit of lower cretaceoussalt diapir 1 2 5 6 3 4 a fig. 13. base cretaceous twt structure map. note the continuation of the børglum fault further to the north than the grenå–helsingborg fault. bf, børglum fault, a and b are strands of the børglum fault; ghf, grenå–helsingborg fault. marine to fluvial gassum formation (nielsen et al. 1989; nielsen 2003, this volume). instead differential subsidence took over in the central part of the kattegat, within the sorgenfrei–tornquist zone. in this area, no mobile zechstein salt is present and subsidence was controlled by basement-attached faults only. no differential subsidence seems to have taken place further to the northwest in the danish part of the northern zechstein salt basin (fig. 7). jurassic in the early jurassic, subsidence controlled by basementattached faults continued within the sorgenfrei–tornquist zone in the areas where no mobile zechstein salt was present. in addition, differential subsidence now also started further to the north-west in areas underlain by mobile zechstein salt (fig. 8). this change indicates increased fault activity (norling & bergström 1987) and in particular salt withdrawal subsidence, possibly triggered by the faulting. activity along nnw–sse-trending normal faults outside the sorgenfrei–tornquist zone (fig. 8) points to a wsw–ene orientation of the minimum stress axis. this again would indicate right-lateral movements along the nw–se-oriented major faults of the sorgenfrei–tornquist zone, as in the triassic. however, since the minimum principal stress axis was at a higher angle to these faults, a reduced amount of strike-slip compared to the triassic may be suggested. the local depocentre in the central part of profile 2 (figs 3, 8), could be regarded as a salt withdrawal basin, similar to those further to the north-west (koyi & petersen 1993; christensen & korstgård 1994). however, no major accumulation of zechstein salt is present in the area, and the immobile marginal zechstein basin facies starts approximately here (mogensen 1994). a section across this depocentre (line a in fig. 8) backstripped (fig. 14a–e) shows no salt structure evolution matching the size of the jurassic – lower cretaceous depocentre (figs 14b, e). the jurassic – lower cretaceous depocentre therefore cannot be explained by salt withdrawal. an explanation could be that the grenå–helsingborg fault seems to terminate in this area (figs 4, 8, 12, 13). termination of a fault having right-lateral transtensional displacement along strike, causes primarily horsetail splaying and sagging (christie-blick & biddle 1985; harding et al. 1985; sylvester 1988). this kind of sagging would cause subsidence that might be only slightly influenced by faulting. the configuration of the jurassic – lower cretaceous depocentre (fig. 14b), is hardly influenced by seismically resolvable faults, and it is therefore proposed that this particular depocentre was caused by dextral transtensional sagging at the north-westernmost termination of the grenå–helsingborg fault. during the late cretaceous – early tertiary inversion tectonics (dextral transpression as opposed to the former dextral transtension), this depocentre was inverted as a closed anticline (fig. 13; see also mogensen & jensen 1994). deposition changed from dominantly marine in the early jurassic to more shallow marine to continental dominated in the mid-jurassic, possibly due to uplift of the crustal block to the north-east of the sorgenfrei– tornquist zone. this change is also reflected in the difference in the depositional pattern between the lower jurassic fjerritslev formation (fig. 8) and the middle jurassic haldager sand formation (fig. 9), notably the change in thickness. fault-related deposition seems to have been limited in the middle jurassic, although faulting along the major faults of the sorgenfrei–tornquist zone created space for middle jurassic deposits in the central part of the kattegat (fig. 9), and caused volcanic activity in skåne, mainly along nw–se-trending faults and fracture zones (erlström et al. 1997). in the central kattegat area and skåne, only limited deposition took place outside the sorgenfrei–tornquist zone during the middle jurassic (fig. 9; norling & bergström 1987). due to the limited faulting it is difficult to deduce any stress orientations. regional indications such as middle jurassic normal faulting in the north sea central graben (mogensen et al. 1992) suggest an e–w to ese–wsw orientation, and if this orientation is extended to the kattegat area, right-lateral movements would again have been induced along the major faults of the sorgenfrei–tornquist zone. inside the sorgenfrei–tornquist zone, deposition may have been continuous from the triassic to the late jurassic (michelsen & nielsen 1991; seidenkrantz et al. 1993), in contrast to adjacent areas outside the zone (michelsen 1986, 1989; nielsen 2003, this volume). during the late jurassic, new marine transgressions invaded the area, filling former irregular topography in skåne (norling & bergström 1987). the early cretaceous had the same depositional evolution, except that tectonic activity increased (ziegler 1987, 1990; eugeno-s working group 1988). differential subsidence took place along the sorgenfrei–tornquist zone in the central kattegat area (fig. 10) and in skåne (norling & bergström 1987) during both the late jurassic and the early cretaceous. large parts of the upper jurassic – lower cretaceous deposits were removed by subsequent erosion, obscuring the depositional pattern in 452 the kattegat area (figs 10, 11). however, thickening of the upper jurassic and lower cretaceous successions towards the eroded area and the børglum fault (fig. 3, profiles 3 and 5), indicates that this was the most active fault in the central kattegat area during both the late jurassic and the early cretaceous, as well as during the late cretaceous – early tertiary inversion phase (figs 10, 11). the orientation of a small upper jurassic depocentre just west of the terne-1 well, related to a n–s-trending fault (fig. 10), might indicate that e–w extension in the kattegat area also persisted during the late jurassic. this is in agreement with a proposed regional late jurassic e–w extension, as in the north sea area (bartholomew et al. 1993; sears et al. 1993). such an orientation of the tensional stresses would again cause right-lateral transtension in the kattegat area along the nw–se-oriented major faults of the sorgenfrei–tornquist zone. right-lateral transtension might also be indicated by the continuous evolution of the local depocentre at line a (figs 10, 11) and profile 2 (fig. 3) in late jurassic (fig. 10) and in early cretaceous times (figs 10, 11). this development could also have been caused by sagging due to dextral transtensional fault termination, as proposed previously. indications of the sense of lateral transtensional displacement along the sorgenfrei–tornquist zone boundary faults become more obscure in mid-jurassic – early cretaceous times, due to the effect and overprint of later inversion tectonics. the few clear indications of the sense of lateral displacement in this period as well as in all former mesozoic periods, seem to favour rightlateral displacement. we therefore suggest a general dextral transtensional displacement along the nw–setrending grenå–helsingborg fault, and in particular the børglum fault, during the whole of the mesozoic (for 453 present base chalk top ‘gassum fm’ near top lower triassic near top lower triassic, top pre-zechstein is flattenede d c b a ssw nne grenå–helsingborg fault u. cretaceous l. jurassic – l. cret. u. triassic – l. jurassic (gassum fm) l. triassic – u. triassic rotliegendes syn-rift sequence l. palaeozoic precambrian basement fig. 14. backstripping sequence across the grenå–helsingborg fault (fig. 2, line a; seismic line k84-001). the anticlinal structure in the lower cretaceous succession was created by inversion of a late triassic to early cretaceous syncline. the configuration of this depocentre (crosses on profile b) could indicate salt withdrawal subsidence over an escaping salt pillow, but backstripping down to the surface of ‘no differential subsidence over fault’, close to top lower triassic (profile d), gives no indication of salt structure development (hatched area on e) that could match the depocentre on b. another explanation could therefore be right-lateral transtensional sagging at the tip of the grenå–helsingborg fault which terminates in this area. 454 displacement during the late cretaceous, see mogensen & jensen 1994). this contrasts with the view of pegrum (1984) who proposed left-lateral displacement, especially during the latter part of the mesozoic. an interpretation of the fault configuration in the south-eastern part of kattegat by aubert (1988, fig. 56) also favoured leftlateral displacement. our interpretation of a more closelyspaced seismic grid has changed this fault configuration into a major bend in the grenå–helsingborg fault (fig. 4). right-lateral movements along the grenå– helsingborg fault start to be transferred to the børglum fault, in order to accommodate this major bend. sivhed (1991) suggested 4 km of left-lateral displacement along the fyledal fault, skåne, based on the displacement of a possible post-early permian channel. however, age relationships between channels on both sides of the fyledal fault appear to be uncertain because of different lithologies in the channel fill (sivhed 1991). left-lateral movement would be possible on a fault in terne-1 hans-1 sæby-1 frederikshavn-(1–3) skagen-2 rønde-1 lavø-1 anholt-(1–4) voldum-1 20 km sweden denmark ghf bf 57°n 56°n 12°e11°e fig. 16. major palaeozoic faults of the kattegat area, which, when compared with the different mesozoic maps (figs 4–13), show that almost all mesozoic faults are reactivated palaeozoic faults. bf, børglum fault; ghf, grenå–helsingborg fault. grenå–helsingborg fault børglum fault triassic – early cretaceous dextral transtensional evolution late cretaceous – early tertiary dextral transpressional evolution the pre-mesozoic sorgenfrei–tornquist zone fig. 15. schematic diagram showing the response of the sorgenfrei–tornquist zone to a changing stress field and illustrating how the zone acted as a buffer zone between more coherent crustal blocks, whenever changes in the regional stress field were induced. 455 triassic regional subsidence, dextral transtension w e w e w e w e w e w e a b c d e f w e g quaternary u. cretaceous m./u. jurassic – l. cretaceous l. jurassic m.–u. triassic l. triassic zechstein (salt/marginal facies) u. carbon./rotliegendes syn-rift sequence u. silurian – devonian cambrian – l. silurian basement lateral movements away from viewer lateral movements towards viewer early palaeozoic foreland deep to the caledonian orogeny? jurassic – early cretaceous extension, minor lateral movements late carboniferous – early permian rift phase, dextral transtension, peneplanisation late cretaceous – early tertiary inversion/dextral transpression neogene regional uplift and erosion zechstein limited deposition in the area fig. 17. summary of phanerozoic tectonic events in the kattegat area and along the sorgenfrei–tornquist zone. a right-lateral fault zone, depending on the orientation of the fault. however, 4 km of sinistral movement along a fault with a nw–se orientation such as the fyledal fault seems unlikely. the sorgenfrei–tornquist zone as a buffer zone we consider the scandinavian part of the tornquist zone to be a very old and weak intercratonic boundary (pegrum 1994), between crustal blocks with thicknesses along the zone in the order of 25–30 km (ro et al. 1990b), and with total lateral displacements of only 20–30 km (mogensen 1994). as a boundary between crustal blocks, the sorgenfrei–tornquist zone is also a deep-seated fault zone where changes in the stress field are more easily accommodated than in the more coherent adjacent crustal blocks (fig. 15). mesozoic reactivation of old, weak basement lineaments in an extensional regime with an oblique angle to the principal stress directions has been described from the north sea (bartholomew et al. 1993; sears et al. 1993). comparison of the faults from the different mesozoic maps (figs 4–13) with the palaeozoic fault pattern (fig. 16) shows that almost all the mesozoic faults are reactivated palaeozoic faults. the weak crustal sorgenfrei–tornquist zone is bounded by the grenå– helsingborg fault and the børglum fault. depending on the orientation of the stress field, these boundary faults may exhibit transpressional or transtensional strike-slip motion. only very special orientations of the stress field will give pure strike-slip motions or pure extension during reactivation. along the børglum fault, a total estimate of 7 km of right-lateral displacement, from the late palaeozoic to the present, was proposed by mogensen (1994). quantification of the amount of lateral displacement during different tectonic episodes is difficult, and can only be done relative to each episode. the period with the largest dextral movements was probably the triassic, as estimated from the configuration of small fault-bend related depocentres and the regional stress orientation (see above). as much as half of the dextral displacement (3–4 km) could be attributed to the triassic movements. this leaves 3–4 km to the rest of the mesozoic transtensional and transpressional movements, of which the late cretaceous – early tertiary transpressional movements were probably the most important. dextral displacement during the jurassic – early cretaceous therefore must have been only a few kilometres. conclusions a better understanding of the triassic–jurassic (–early cretaceous) tectonic processes along the sorgenfrei– tornquist strike-slip fault zone has been provided through the interpretation of closely-spaced 2d reflection seismic data in the kattegat area, denmark. based on this interpretation, several maps have been generated, which in combination with key regional seismic sections outline in detail the structural development of the kattegat area and the sorgenfrei–tornquist zone during this period (fig. 17). during the permian, the area was exposed to erosion and was peneplaned. regional triassic subsidence and tilting resulted in onlap towards the north-east, where the youngest triassic sediments are found overlying precambrian crystalline basement. during the early triassic, in particular, several of the major late carboniferous – early permian faults were reactivated, with dextral strike-slip along the børglum fault. differential subsidence within the sorgenfrei–tornquist zone started at the transition between the late triassic and the early jurassic, primarily with deposition of the fjerritslev formation. this differential subsidence was restricted mainly to the area between the two main faults in the sorgenfrei–tornquist zone, the grenå–helsingborg fault and the børglum fault. the restricted basin development indicates a change in the regional stress field. subsidence during the middle jurassic and the late jurassic – early cretaceous followed the early jurassic pattern with differential deposition within the sorgenfrei–tornquist zone, but now even more restricted to the zone. the early cretaceous subsidence pattern was a direct continuation of the late jurassic subsidence with no hiatus in between. the only difference was the increased rate of subsidence during early cretaceous times. many small faults were generated during the mesozoic in the area between the terne-1 and hans-1 wells and the grenå–helsingborg and børglum faults. this fault pattern indicates a general transfer of strike-slip/obliqueslip motion from the grenå–helsingborg fault to the børglum fault. reactivation of old basement faults caused dextral movements along the major boundary faults of the nw–se-oriented sorgenfrei–tornquist zone during the entire mesozoic due to the orientation of the regional stress field. 456 457 acknowledgements we thank s. olausson and o. simonsen for comments on the manuscript. o.r. clausen and j.e. christensen contributed to our discussions during preparation. we also thank the reviewers, r.m. pegrum and o.v. vejbæk for constructive comments and f. surlyk for numerous suggestions for improvements. t.e. mogensen was supported by a grant from the danish natural science research council. references aubert, k. 1988: strukturell og stratigrafisk utvikling i kattegat, 126 pp. unpublished cand. scient. thesis, oslo universitet, norge. aydin, a. & nur, a. 1985: the types and role of stepovers in strikeslip tectonics. in: christie-blick, n. & biddle k.t. 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(ed.): rift zones in the continental crust of europe – geophysical, geological and geochemical evidence: oslo–horn graben. tectonophysics 178, 29–49. ziegler, p.a. 1987: late cretaceous and cenozoic intra-plate compressional deformations in the alpine foreland – a geodynamic model. in: ziegler, p.a. (ed.): compressional intra-plate deformations in the alpine foreland. tectonophysics 137, 389–420. ziegler, p.a. 1990: geological atlas of western and central europe, 2nd edition, 239 pp. amsterdam: elsevier for shell internationale petroleum maatschappij. manuscript received 1 november 1995; revision accepted 2 september 1997. 16_jeb_cet.indd 53 albedo, latin for ‘whiteness’, is a term used to describe the amount of sunlight reflected by the ground. fresh snow albedo can exceed 85%, making it among the most reflective natural substances. warm conditions promote snow crystal metamorphosis that, like the presence of liquid water, bring snow albedo down below 65%. with the darkening, caused by the metamorphosis, absorbed solar energy thus increases by roughly a factor of two. seasonal snow melts over the lower reaches of a glacier leading to the exposure of bare ice with albedo below 55%. impurities such as dust, black carbon or microbes can bring glacier-ice albedo below 30%, meaning that snow ablation gives way to impurity-rich, bare glacier ice which increases absorbed sunlight by more than a factor of three. the thickness of the winter snow layer and the intensity of spring melt are important determinants of the annual glacier-ice melt, as the amount of snow cover governs the timing of darker ice exposure; the earlier the exposure, the more ice can melt. because snow and ice albedo properties make it an amplifier of climate change, surface albedo has been designated as an essential climate variable and a target requirement for climate monitoring (wmo 2011). polar orbiting satellites facilitate albedo mapping with arctic coverage multiple times per day in clear-sky conditions. satellite-based retrievals of surface albedo depend on accurate compensation of the intervening atmosphere. thus, without ground truth, the satellite retrievals are uncertain. in greenland, snow and ice albedo is monitored by automatic weather stations (awss) from the greenland climate network (gc-net; steffen et al. 1996) since 1995 and after 2007 from the programme for monitoring of the greenland ice sheet (promice; van as et al. 2013). using the gc-net data, satellite-derived albedo values are compared with ground data (e.g. stroeve et al. 2013). here, we present comparisons of daily gc-net and promice albedo data to satellite-derived albedo from the nasa moderate resolution imaging spectroradiometer (modis) mod10a1 product (hall et al. 1995). mod10a1 data have been available since may 2000 and are here de-noised, gap-filled and calibrated into a daily 500 × 500 m grid covering greenland, iceland and the canadian arctic glaciers (fig. 1). daily albedo from modis the mod10a1 product contains daily snow extent, snow albedo, fractional snow cover and a data quality assessment at 500 × 500 m resolution (hall et al. 1995). both nasa terra and aqua satellites are equipped with modis sensors. here, terra data are chosen over aqua data as they greenland, canadian and icelandic land-ice albedo grids (2000–2016) jason e. box, dirk van as, konrad steffen and the promice project team* * robert s. fausto, andreas p. ahlstrøm, michele citterio and signe b. andersen fig. 1. an example (15 july 2016) of the daily 500 m × 500 m enhanced mod10a1 collection 6 albedo product presented here. the circles show positions of ground-truth observations. 84 80 76 72 68 64 60 56 52 48 44 40 36 32 28 albedo (%) promice gc-net kcp jar © 2017 geus. geological survey of denmark and greenland bulletin 38, 53–56. open access: www.geus.dk/publications/bull 5454 give longer temporal coverage, and aqua modis band 6 detectors (useful in cloud discrimination) have become degraded or non-functional (modis characterization support team, nasa, updated february 2017). an alternative modis albedo product (mcd43) was not chosen due to its reduced temporal resolution of eight days. the mod10a1 data used here span the arctic melt season 15 march (74th day of year) to 26 october (299th day of year) for the 17 year period 2000 to 2016. the two latest mod10a1 versions are evaluated; collection 5 (hall et al. 2011) and collection 6 (riggs & hall 2015; hall & riggs 2016), hereafter c5 and c6. mod10a1 de-noising, smoothing and gap-filling inspection of the c5 and c6 albedo imagery reveals that, despite some cases when pixel quality is coded ‘best’ or ‘good’, cloud artifacts resembling shadows, aircraft condensation trails, thin clouds, and cloud edges can persist, often over the brightest areas presumably where there is less distinction between clouds and clean snow. fortunately, because the artifacts introduce abrupt temporal departures in the albedo (α) time series, it is possible to reject them on a pixel by pixel basis using temporal statistics from multi-day albedo (αn-days) samples. here, an 11 day αn-days sample size is selected; five days before and after each day i. on a pixel-by-pixel basis, statistics are computed from αn-days. the number of days n does not always represent 11 albedo values because some days a pixel may already be dismissed as cloudy, missing or of inadequate quality. only cases with at least four samples per 11-day window are considered sufficient for an albedo estimate for that day and pixel. the final pixel by pixel daily albedo values are taken as the 11-day average of available values below a fractional noise threshold (d) value of 0.4, with d computed as: di = | (αn-days median(αn-days)) / median(αn-days) | for low albedo variability areas, for example the dry snow area, when the standard deviation αn-days is under 0.03, then a more strict d threshold of 0.1 is used. the procedure has both a smoothing and a gap filling effect on the albedo time series. the resulting data product can be viewed at https://tinyurl.com/promice-albedo-greenland. modis validation using ground-truth albedo daily promice and gc-net aws albedo values are compared with the nearest 500 × 500 m resolution mod10a1 c5 and c6 values for all stations in each of 9 or 16 years that span 2007 to 2016 or 2000 to 2015, respectively. figure 2 provides an example for c6 data illustrating a typical result of the de-noising procedure, yielding an increased number of modis values, increased correlation, reduced root mean squared difference (rmsd) shifting regression slope closer to unity and no real change in bias. table 1 lists summary statistics for the multi-year, multistation comparison with mod10a1 c5 and c6. the mod10a1 skill either improves or is stable in the c5 to c6 update. the average bias and root mean squared difference decrease and the correlation and average count of days increase. the number of compared station-years increases. from raw to de-noised, there is also a consistent improvement in agreement between the satellite and ground data (table 2). in the de-noised product, the rmsds are 0.08 for promice stations that are concentrated in the ablation day of year al be do (% ) 100 150 200 250 90 80 70 60 50 40 de-noised mod10a1 promice n = 117 regression: correlation = 0.962 promice = modis*0.872 + 0.090 bias = 0.008 rmsd = 0.043 pr o m ic e alb ed o (% ) mod10a albedo (%) 40 50 60 70 80 90 90 80 70 60 50 40 fig. 2. year 2013 example comparison of daily de-noised albedo from satellite (nasa modis mod10a1 collection 6 data) and the ground (promice) for the kpc_l (fig. 1) station on the north-eastern greenland ice sheet. 55 area or 0.05 for gc-net stations that are concentrated in the accumulation area. e lower gc-net correlation and rmsd values result from the mostly dry snow areas where albedo variability is small. for promice stations which are concentrated in the ablation area (and for the gc-net jakobshavn ablation region (jar) stations), the larger mod10a1 pixel footprint includes a complex contribution from some combination of e.g., crevasses, snow patches and concentrated or distributed snow and ice impurities such as cryoconite. e root mean squared di erence is probably more attributable to the ground data because they have a c. 5–10 m2 footprint, four orders of magnitude smaller than the mod10a1 500 m × 500 m footprint. modis validation using gc-net albedo gc-net albedo data, having a time coverage longer than 10 years, are compared with c5 and c6 to evaluate accuracy in year-to-year albedo changes. mod10a1 collection 6 (hall & riggs 2016) compensates modis sensor degradation found in collection 5 (lyapustin et al. 2014). c6 is found to compensate the temporal trend bias in dry snow areas (fig. 3a). e trend bias is usually smaller or nonexistent for darker targets such as the ablation area (fig. 3b). e 0.02 albedo o set at the summit site is partly attributable to the bias described in the following. mod10a1 sun angle bias whereas the adjustments to collection 6 eliminate a spurious darkening trend concentrated over snow and in the northern part of greenland (polashenski et al. 2015), both collections 5 and 6 mod10a1 albedo products have a residual bias based on the angle of the sun above the horizon. e bias is evident over nearly 20° of latitude range of the promice and gc-net data. in april (days 91–120), there is no bias in southern greenland but a c. 4% bright bias in the northern 2/3 of greenland (fig. 4a). by june (days 152–181), the pattern of the bias has shifted to a more uniform dark bias strongest in the south (fig. 4b). e bias varies over time and latitude (see the animation: https://ti nyurl.com/bias-vs-lat). we correct the collection 6 bias according to the daily variation in the regression line (blue dashed line in fig. 4a, trends in fig. 4b). e calibration table 1. summary statistics for comparison of αmod10a1 collection 5 and 6 with available αpromice and αgc-net in the 2000–2016 period (de-noised). collection 5 correlation 0.838 0.182 bias –0.024 0.035 rmsd * 0.086 0.042 a verage count of days 117 37 collection 6 correlation 0.832 0.183 bias –0.002 0.034 rmsd 0.084 0.044 a verage count of days 124 39 collection 5 correlation 0.490 0.225 bias -0.006 0.099 rmsd 0.104 0.090 a verage count of days 114 33 collection 6 correlation 0.581 0.259 bias 0.006 0.038 rmsd 0.050 0.033 a verage count of days 110 32 αmod10a1 versus αpromice 144 station years average st. dev. αmod10a1 versus αgc-net 183 station-years average st. dev. * root mean squared difference. table 2. summary statistics for comparison of αmod10a1 collection 5 and 6 with available αpromice and αgc-net in the 2000–2016 period (de-noised minus raw). collection 5 correlation 0.067 0.003 bias –0.001 0.001 rmsd * –0.022 0.003 count of days per year 28 13 collection 6 correlation 0.070 0.032 bias –0.002 0.001 rmsd –0.029 0.003 count of days per year 28 13 collection 5 correlation 0.111 –0.048 bias 0.002 0.026 rmsd 0.016 0.031 count of days per year 32.000 11.000 collection 6 correlation 0.170 –0.005 bias –0.005 –0.042 rmsd –0.030 –0.037 count of days per year 29 12 αmod10a1 versus αpromice 4 station years average st. dev. αmod10a1 versus αgc-net 19 station-years average st. dev. * root mean squared difference. year year al be do (% ) al be do (% ) 2000 2005 2010 2015 2000 2005 2010 2015 summit (72.5°n, 38.5°w, 3254 m a.s.l.) gc-net trend: –1.0% mod10a1 collection 5 trend: –5.4% mod10a1 collection 6 trend: –1.2% 84 83 82 81 80 79 70 60 50 40 30 20 jar1 (69.4°n, 49.6°w, 962 m a.s.l.) gc-net trend: –27.0% mod10a1 collection 5 trend: –27.0% mod10a1 collection 6 trend: –22.9% ba fig. 3. examples of july monthly average dry snow area (a) and ablation area (b) modis collection 5 and 6 compared with gc-net albedo trends spanning 16 summers. jar: jakobshavn ablation region. nr38_p53-56.pdf 1 09/08/17 09.18 nr38_p53-56.pdf 1 09/08/17 09.36 https://tinyurl.com/bias-vs-lat 5656 assumes there is no sun-angle-dependent bias in the promice and gc-net data. modis albedo for iceland and the canadian arctic the regional product also includes albedo for glaciated areas in iceland, jan mayen and the canadian arctic. the occurrence of clouds reduces the coverage of the product, in the case of land ice, especially at the lowest elevations often near oceans. iceland has data coverage less than half of that of most areas of greenland. areas of the canadian arctic such as the devon ice cap also have reduced time coverage compared to greenland, which has a stronger cloud clearing effect from its high pressure areas often centered over the ice sheet. a new promice data product the methodology developed here for de-noising, gap filling, and bias correction for the mod10a1 albedo product yields an enhanced modis mod10a1 climate data record available for download through the promice database via the webpage www.promice.dk acknowledgements promice is supported by the danish energy agency through the dancea programme. gc-net is supported by nasa and us national science foundation grants. references hall, d.k. & riggs, g.a. 2016: modis/terra snow cover daily l3 global 500m grid, version 6. greenland coverage. national snow and ice data center, nasa distributed active archive center, boulder, colorado usa. http://nsidc.org/data/mod10a1/versions/6, accessed december 2016. hall, d.k., riggs, g.a. & salomonson, v.v. 1995: development of methods for mapping global snow cover using moderate resolution imaging spectroradiometer data. remote sensing of environment 54, 127–140. hall, d.k., riggs, g.a. & salomonson, v.v. 2011: modis/terra snow cover daily l3 global 500 m grid v004, january to march 2003. digital media, updated daily. national snow and ice data center, boulder, co, usa. lyapustin, a. et al. 2014: science impact of modis c5 calibration degradation and c6+ improvements. atmospheric measurement techniques discussion 7, 7281–7319. modis characterization support team, nasa, updated february 2017: http://mcst.gsfc.nasa.gov/calibration/time-dependent-list-nonfunctional-or-noisy-detector polashenski, c.m., dibb, j.e., flanner, m.g., chen, j.y., courville, z.r., lai, a.m., schauer, j.j., shafer, m.m. & bergin, m. 2015: neither dust nor black carbon causing apparent albedo decline in greenland’s dry snow zone: implications for modis c5 surface reflectance, geophysical research letters 42, 9319–9327. riggs, g.a. & hall, d.k. 2015: modis snow products collection 6, user guide. https://nsidc.org/sites/nsidc.org/files/files/modis-snowuser-guide-c6.pdf steffen, k., box, j.e. & abdalati, w. 1996: greenland climate network: gc-net. in: colbeck, s.c. (ed.): crrel 96-27 special report on glaciers, ice sheets and volcanoes, tribute to m. meier, 98– 103. hannover: u.s. army. stroeve, j.c., box, j.e., wang, z., schaaf, c. & barrett, a. 2013: reevaluation of modis mcd43 greenland albedo accuracy and trends. remote sensing of environment 138, 199–214. van as, d., fausto, r.s., colgan, w.t., box, j.e. and the promice project team 2013: darkening of the greenland ice sheet due to the melt-albedo feedback observed at promice weather stations. geological survey of denmark and greenland bulletin 28, 69–72. wmo (world meteorological organization) 2011: systematic observation requirements for satellite-based data products for climate, update. global climate observing system, gcos-154, 138 pp. latitude (°n) day of year al be do b ias (u ni tle ss ) 65 70 75 80 100 120 140 160 180 200 220 240 promice gc-net 0.06 0.04 0.02 0.00 –0.20 –0.04 –0.06 0.08 0.06 0.04 0.02 0.00 –0.20 –0.04 al be do b ias (u ni tle ss ) 60°n 70°n 80°n a b fig. 4. a: example of nasa modis mod10a1 collection 6 difference with ground data from automatic weather stations versus latitude, indicating a sun-angle-dependent bias. b: bias for three latitude bands. the station names are abbreviated. authors’ addresses j.e.b., d.v.a. & the promice team, geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark; e-mail: jeb@geus.dk k.s., swiss federal institute for forest, snow and landscape research, zürcherstrasse 111, ch-8903 birmensdorf, switzerland. denoised mod10a1 promice: n 117 regression correlation 0962 promice modis0872 0090 bias 0008 rmsd 0043: 60n 70n 80n b: undefined: undefined_2: authors addresses: geological survey of denmark and greenland bulletin 4, 2003, pp 21-24 the search for new, deep-seated drinking water resources in denmark has increased significantly during the past five years as a result of the discovery of excessive amounts of nitrate, pesticides and other pollutants in shallow groundwater boreholes (e.g. nygaard et al. 2004, this volume). to find and map these aquifers, a multidisciplinary sequence stratigraphic approach has successfully been applied to the miocene deposits of southern jutland, where especially the odderup and ribe formations are known as a main aquifer for drinking water from several test wells (rasmussen et al. 2002). recently, a more systematic study of the miocene succession in central and western jutland has been initiated by the geological survey of denmark and greenland (geus) under contract with local authorities. it includes detailed sedimentological descriptions of outcrops, sedimentological and log-interpretations of new stratigraphic boreholes and interpretation of new high-resolution seismic data (fig. 1). a number of outcrops and wells have been studied palynologically, resulting in a detailed dinoflagellate cyst stratigraphy and in palynofacies interpretations. the results of these studies have been integrated in the regional geological and stratigraphic model (fig. 2). two new aquifers have been discovered: the bastrup sand and the billund sand. the bastrup sand has already been exploited as a main aquifer in central and southern jutland, and has been referred to either the ribe or odderup formations. however, new stratigraphic results reveal that the bastrup sand is a separate unit in the miocene succession. the billund sand is a deep-seated aquifer located more than 100 m and often more than 150 m deep, and is therefore not penetrated by standard water supply wells which rarely reach c. 100 m. the billund sand was first revealed by multichannel seismic data deriving from former oil-exploration carried out in 21 the billund delta: a possible new giant aquifer in central and western jutland erik s. rasmussen, karen dybkjær and stefan piasecki fig. 1. map of southern jutland showing palaeoenvironments of the billund delta, the location of boreholes and the seismic section mentioned in the text. fig. 2. lithostratigraphy of the miocene succession in jutland compiled from larsen & dinesen (1959), rasmussen (1995), dybkjær & rasmussen (2000) and rasmussen et al. (2002). bs, billund sand; gs, gram silt/sand; hs, hvidbjerg sand; vfs, vejle fjord sand member. geological survey of denmark and greenland bulletin 4, 21–24 (2004) © geus, 2004 the billund area (fig. 3a). the resolution of these seismic data is very poor, but one interpretation of the dipping reflectors (clinoforms) seen in fig. 3a was of a delta complex. this agrees with outcrop studies along the fjords of eastern jutland which suggest that a spit complex was deposited in this area during the early miocene. the billund sand was tested by the vandel mark well in 2001, which penetrated c. 40 m of sand at a depth of 200 m. the presence of a regional major sand body was later confirmed by new high-resolution seismic data and by the billund and løvlund wells in 2002. the billund well penetrated 50 m of mediumto coarse-grained sand, and chemical tests of the water quality were good. however, a water supply well at fjand in western jutland has had problems with so-called ‘brown water’ – water enriched in organic matter (humus). saline water may also be expected close to older deep-seated faults. this paper summarises the results of a mapping programme of the billund sand initiated in the summer of 2003. 22 fig. 3. two seismic sections from the billund area. a: the old multichannel seismic line dcj-05. b: the new shallow seismic line gi01 (courtesy of cowi a/s). the red framed area in fig. 3a corresponds to the section shown in fig. 3b. the clinoformal reflection pattern is indicated in yellow. fig. 4. correlation panel of four boreholes, with gamma-ray log signature, trending n–s in the billund area. the seismic section adjacent to the store vorslunde well shows the correlation of a clinoformal reflection pattern with sand. note that the billund delta pinches out between the vandel mark and the almstok wells. seismic data courtesy of cowi a/s. the billund delta the billund sand represents a major early miocene delta prograding southwards from norway into the danish area. the study of the billund delta includes interpretation of seismic data, lithological descriptions and interpretation of logs from new boreholes, sedimentological descriptions and interpretation of outcrops, dating and correlation of the succession by biostratigraphy (mainly dinoflagellates) and finally interpretation of the depositional environment based on palynofacies. integration of these disciplines has resulted in a robust geological model for the delta that provides a basis for the prediction and location of good reservoir rocks suitable as aquifers for drinking water. seismic mapping in order to map the extent and thickness of the delta, seismic data from both the north sea and jutland have been used. based upon the study of the seismic data an isochore map of the delta has been constructed. the delta extends from 50 km west of the present-day west coast of jutland trending nw–se across jutland, from ringkøbing in the west to billund in central jutland. east of billund the main delta bends northwards, and the eastern limit follows a line from the town of give northwards. the estimated size of the delta is in the order of 10 000 km2 whereas its thickness varies from 300 m in the north sea to less than 100 m in central jutland. associated spit complexes that outcrop in eastern jutland are rarely thicker than 30 m. in some areas the seismic data are characterised by strongly southwards-dipping reflectors (clinoforms; fig. 3b). correlation of seismic data with borehole information reveals that where the clinoforms are distinct, they represent massive sands (fig. 4); north of billund, the massive sands may be up to 75 m thick. the billund delta complex is characterised by a pinchout distance of c. 2 km which in recent delta systems is a characteristic feature of wave-dominated deltas (cf. løseth & helland-hansen 2001). study of boreholes a number of recent, deep boreholes have penetrated the billund delta (fig. 4). these show that the delta comprises a coarsening-upwards succession of mediumto coarse-grained sand with gravel and thin clay layers locally present. a very coarse-grained layer succeeded by a thinning-upwards succession in the store vorslunde borehole indicates that some of the sand was deposited in channels. the thickness of the delta sand varies from 40 m in store vorslunde to 50 m in the billund borehole. however, the eg-3 well (fig. 1), an old deep exploration well, indicates that up to 100 m of sand may be present. study of outcrops a spit system corresponding to the down-drift part of the billund delta outcrops along the fjords in eastern jutland (fig. 5; friis et al. 1998). these deposits are known as the vejle fjord formation (fig. 2). during the last five years these outcrops have been relatively well exposed, and new sequence stratigraphic, biostratigraphic and sedimentological studies have been carried out indicating that a major nw–se-trending spit was located near the southern part of vejle fjord. north of the spit, organic-rich, fine-grained sediments were laid down in a lagoon. sand-rich lagoonal sediments were deposited near inlets, e.g. at dykær. south of the 23 fig. 5. block diagram showing the depositional model of the billund delta. in the upper block diagram the seismic panel shows a prograding delta and the pinchout between the billund and almstok wells. the lower diagram illustrates data from outcrops and wells indicating the presence of a spit complex. note that the spit complex correlates with the delta front at store vorslunde. seismic data courtesy of cowi a/s. 24 spit, ebb-dominated tidal deltas were deposited; examples of these are exposed in a sandpit at pjedsted (fig. 6). at lillebælt, upper and lower shoreface sediments were deposited, and are excellently exposed in coastal cliffs at the old lillebælt bridge (lillebæltsbro) on funen and at børup (fig. 1). the sedimentological study of the vejle fjord formation indicates that the sediments were laid down in a wave-dominated depositional environment with some tidal influence (rasmussen et al. 2002). future perspectives the billund delta may be one of the most important future aquifers for drinking water in western and central jutland. furthermore, the excellent seismic data and the good outcrops in eastern jutland make it possible to use the billund delta as a good analogue for jurassic reservoir rocks in the central and viking grabens of the north sea. references dybkjær, k. & rasmussen, e.s. 2000: palynological dating of the oligocene–miocene successions in the lille bælt area, denmark. bulletin of the geological society of denmark 47, 87–103. friis, h., mikkelsen, j. & sandersen, p. 1998: depositional environment of the vejle fjord formation of the upper oligocene – lower miocene of denmark: a back island/barrier-protected depositional complex. sedimentary geology 17, 221–244. larsen, g. & dinesen, a. 1959: vejle fjord formation ved brejning: sedimenterne og foraminiferfaunaen (oligocæn–miocæn). danmarks geologiske undersøgelse ii. række 82, 114 pp. løseth, t.m. & helland-hansen, w. 2001: predicting the pinchout distance of shoreline tongues. terra nova 13, 241–248. nygaard, e., ernstsen, v., jacobsen, c.s., jacobsen, o.h., juhler, r.k., van der keur, p., olesen, s.e., rasmussen, j., rosenberg, p. & vosgerau, h. 2004: pesticide leaching in danish groundwater: identification of vulnerable areas. geological survey of denmark and greenland bulletin 4, 25–28 (this volume). rasmussen, e.s. 1995: vejle fjord formation: clay mineralogy and geochemistry. bulletin of the geological society of denmark 42, 57–67. rasmussen, e.s., dybkjær, k. & piasecki, s. 2002: miocene depositional systems of the eastern north sea basin, denmark. development of sedimentological and stratigraphical principles in modern sedimentology. danmarks og grønlands geologiske undersøgelse rapport 2002/89, 131 pp. fig. 6. tidal sandstone deposits from an ebb-dominated delta succession associated with a spit complex, deposited down-drift from the main billund delta. pjedsted sandpit, for location see fig. 1. width of track approx. 50 cm. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: esr@geus.dk << /ascii85encodepages false /allowtransparency false /autopositionepsfiles true /autorotatepages /all /binding /left /calgrayprofile (dot gain 20%) /calrgbprofile (srgb iec61966-2.1) /calcmykprofile (u.s. web coated \050swop\051 v2) /srgbprofile (srgb iec61966-2.1) /cannotembedfontpolicy /warning /compatibilitylevel 1.4 /compressobjects /tags /compresspages true /convertimagestoindexed true /passthroughjpegimages true /createjdffile false /createjobticket false /defaultrenderingintent /default /detectblends true /colorconversionstrategy /leavecolorunchanged /dothumbnails false /embedallfonts true /embedjoboptions true /dscreportinglevel 0 /syntheticboldness 1.00 /emitdscwarnings false /endpage -1 /imagememory 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<feff4f7f752890194e9b8a2d5b9a5efa7acb76840020005000440046002065874ef65305542b8f039ad876845f7150cf89e367905ea6ff0c4fbf65bc63d066075217537054c18cea3002005000440046002065874ef653ef4ee54f7f75280020004100630072006f0062006100740020548c002000520065006100640065007200200035002e0030002053ca66f465b07248672c4f86958b555f3002> >> >> setdistillerparams << /hwresolution [2400 2400] /pagesize [595.000 842.000] >> setpagedevice geological survey of denmark and greenland bulletin 28, 2013, 61-64 61 calibration of spectral gamma-ray logs to deltaic sedimentary facies from the cretaceous atane formation, nuussuaq basin, west greenland gunver krarup pedersen, niels h. schovsbo and henrik nøhr-hansen gamma-ray logs are widely used as a lithology indicator in wells as part of standard petrophysical interpretations. in cored wells, gamma-ray logs should always be calibrated to the lithology in order to correct the petrophysical model. gamma radiation is emitted from three elements, k, th and u (potassium, thorium and uranium) which occur in minerals such as feldspar, mica, glauconite, clay minerals, zircon, titanite and apatite as well as in organic complexes. organicrich mudstones usually have high gamma-radiation values and quartz-rich sandstones low values. in many places, upward-coarsening successions are recognisable from the gamma log. the gamma log records the sum of radiation from k, th and u, and their relative contributions are measured in a spectral gamma-ray log. the present case study focuses on spectral gamma-ray characterisation of the deltaic atane formation which shows well-developed, upward-coarsening delta-front deposits in outcrops (fig. 1c). geological setting the nuussuaq basin is a rift basin, which contains the only exposures of cretaceous and paleocene sediments along the west coast of greenland. the siliciclastic sediments are overlain by a thick pile of volcanic rocks (chalmers et al. 1999; dam et al. 2009). during late cretaceous, greenland was characterised by a warm and probably humid climate. the sediments range from alluvial fans overlying deeply weathered precambrian basement through marginally marine deposits to marine deep-water deposits, all referred to the nuussuaq group (dam et al. 2009). the floodplains and delta plains had a rich flora, recorded in well-preserved plant fig. 1. a: geological map of central west greenland showing the location of ataata kuua on the south coast of nuussuaq. b: the atane formation is erosionally overlain by the kangilia formation in the western slope of ataata kuua. the yellow dot marks the drilling site of borehole 247801, d: dyke. height of section c. 500 m. the frame shows the position of fig. 1c. c: the atane formation with depositional environments indicated. note the distinct coarsening-upward successions (triangles). height of section c. 100 m. © 2013 geus. geological survey of denmark and greenland bulletin 28, 61–64. open access: www.geus.dk/publications/bull greenland 70° disko 50 km nuussuaq 53° volcanics ataata kuua precambrian cretaceous sediments atane formationatane formation a c b c kangilia formation 13 0 m d d d 247801 delta front delta plain shoreface channel 60 m 6262 fig. 2. data from borehole 247801: a simplified sedimentological log of the entire core (566 m), a gamma log measured in the borehole to a depth of 320 m, and the new range chart for the dinoflagellate cysts in core samples. api: american petroleum institute units. ggu 247801 is located at 70°19.87´n, 52°55.8´w. the framed interval is shown in fig. 3. the drill site is shown in fig. 1b. d ep th (m ) 25 50 75 100 125 150 175 200 225 250 275 300 325 350 375 400 425 450 475 500 525 550 c hr on o. la te c re ta ce ou s pe rio d/ ep oc h ea rly c on iac ian – e ar ly sa nt on ian ag e sa m pl es 23.70 43.90 54.35 91.50 118.05 138.55 149.70 182.40 219.00 239.60 256.65 290.85 336.20 360.45 391.00 405.30 433.60 464.00 485.70 519.60 545.70 cl eis to sp ha er idi um m ult ifu rc at um isa be lid ini um sp p. od on to ch itin a sp p. ra ph ido din ium fu ca tu m tr ith yro din ium su sp ec tu m ch at an gie lla d itis sim a ch at an gie lla g ra nu life ra ci rc ulo din ium d ist inc tu m h et er os ph ae rid ium d iff ici le la cin iad ini um a rc tic um od on to ch itin a op er cu lat a ol igo sp ha er idi um a lbe rte ns e ol igo sp ha er idi um sp p. pa lae oh ys tri ch op ho ra in fu so rio ide s sp ini din ium c f. ec hin oid eu m sp ini din ium sp p. sp ini fe rit es sp p. ta ny os ph ae rid ium sp p. isa be lid ini um m ag nu m fr om ea fr ag ilis isa be lid ini um co ok so nia e pa lae op er idi niu m p yro ph or um ch at an gie lla sp p. tr ith yro din ium sp p. ex oc ho sp ha er idi um sp p. su rc ulo sp ha er idi um lo ng ifu rc at um h et er os ph ae rid ium h et er ac an th um pa ra lec an iel la ind en ta ta al te rb idi niu m sp p. ch at an gie lla m cin tyr ei se ne ga lin ium sp p. pa lae oc ys to din ium sp p. ? dinoflagellate cysts ru gu biv es cic uli te s r ug os us ru gu biv es icu lat es re du ctu s ap pe nd ici sp or ite s c f. i ns ign is ? ? ? ? ? ? ? ???? ?? ?? ? ?? ? ? ? ? ? ? sp events chatangiella mcintyrei heterosphaeridium difficile palaeoperidinium pyrophorum spinidinium cf. echinoideum, chatangiella ditissima, chatangiella granulifera, heterosphaeridium difficile trithyrodinium suspectum, rugubivesciculites rugosus, isabelidinium spp. 0 100 200 api base top 200 250 300 350 400 450 500 550 tr tr tr tr tr tr 50 100 20 m 150 tr tr tr core from borehole ggu 247801 at an e fo rm at io n q ila ki ts oq m em be r delta front channel delta plain coal shoreface 63 macrofossils and abundant comminuted plant debris. a huge volume of non-marine to shallow marine deposits constitutes the cretaceous atane formation, which is well exposed along 65 km of the south coast of nuussuaq from sea level to altitudes of 500–800 m. seismic data indicate a minimum thickness of 3000 m for the formation (dam et al. 2009). ataata kuua – a narrow fluvial valley at ataata kuua, on the south coast of nuussuaq (fig. 1), shows the deltaic atane formation erosively truncated and overlain by the turbiditic, paleocene kangilia formation (dam et al. 2009). in 1980, the geological survey of greenland drilled a 566 m deep borehole (ggu 247801) at ataata kuua as part of a regional study of the composition and distribution of coal in the atane formation. the entire succession was cored, with 100% recovery, and a gamma log was measured in the upper 320 m of the borehole. this gamma log as well as a simplified sedimentological log of the entire core are shown adjacent to the new biostratigraphical range chart in fig. 2. biostratigraphy – in recent studies of 21 delta-front mudstone samples from core 247801, palynomorphs have been examined (fig. 2). the diversity and density of dinoflagellate cysts, spores and pollen are very low, but the presence of chatangiella granulifera, heterosphaeridium difficile and spinidinium cf. echinoideum in the lower part and chatangiella mcintyrei and spinidinium cf. echinoideum in the middle to upper part indicates an early coniacian age or younger. an early santonian minimum age of the upper part of the core is indicated by the presence of rugubivesciculites spp., the absence of campanian marker species and by the presence of heterosphaeridium difficile, laciniadinium arcticum and spinidinium cf. echinoideum in a sample from the ataata kuua 2004-3 section situated immediately above the site of borehole 247801. the relatively uniform dinoflagellate assemblage recorded throughout the core (fig. 2) supports the interpretation of a relatively high sedimentation rate. sedimentology – the delta deposits of the atane formation represent four depositional environments: delta front, distributary channel, delta plain and shoreface (figs 1c, 2, 3). the delta-front deposits include mudstones, heterolithic sandstones with wave-generated sedimentary structures and wellsorted sandstones, all with comminuted plant debris. the distributary channel deposits are mostly cross-bedded, mediumto coarse-grained sandstones with some feldspar grains. the delta-plain mudstones are interbedded with coal beds or thin sandstone beds. the thin shoreface sandstones contain abundant marine trace fossils and overlie erosive surfaces. the delta-front deposits and the overlying fluvial sandstones form distinctly upward-coarsening units (figs 1c, 2, 3), interpreted as formed during delta progradation. the shoreface sandstones are interpreted as deposited during a transgression. a detailed log of the core is shown in dam et al. 2009 (fig. 43). the dominant minerals in the mudstones are quartz and kaolinite, neither of which contain more than traces of k. small amounts of k-feldspar and mica result in a k content about 1.3–1.8% k2o, significantly lower than the 2.7% k2o of average mudstones (rider 1990). all mudstones and many sandstones in the atane formation contain comminuted debris from higher land plants. total organic carbon (toc) values of the non-marine delta-plain mudstones range from 3 to 15% and include thin coal beds with 50–65% toc (pedersen et al. 2006). the marine delta-front mudstones contain 6–14% toc, with the highest values in the finegrained, lower part, which includes a flooding surface and had a low sedimentation rate. despite this, marine organic particles, such as dinoflagellate cysts, only form a small part compared to terrestrial organic matter. the gamma log obtained in the borehole shows, as expected, low values for the fluvial sandstones whereas the mud-dominated delta-front and delta-plain deposits are difficult to identify from the gamma log alone (fig. 2). in order to document the contributions of k, th and u to the total gamma radiation we measured the spectral gamma radiation (fig. 3). k is mainly located in feldspar, mica and glauconite; th and u are hosted in zircon, titanite, and apatite. clay minerals may contain small amounts of th, and organic complex compounds may contain u. spectral gamma-ray characterisation core scanning – the core interval was scanned at the core laboratory at the geological survey of denmark and greenland using a set-up which allows simultaneous spectral gamma-ray and density measurements. the spectral gamma-ray analysis is carried out using two 15 cm nai (tl) crystals and the bulk density is determined using a caesium source. the scanning speed was 1 cm/min., corresponding to a vertical resolution of approximately 2 cm for the density log. the scanning data thus supply high resolution data to support sedimentary and geochemical data from the core, as exemplified in fig. 3. results – the upward-coarsening successions, which are clearly seen in the field and in the core (figs 1c, 2, 3), are difficult to identify on the total readings of the spectral gamma log (fig. 3). the grain-size trends are, however, reflected in the th and k logs, which are negatively correlated. the mudstones have high th and low k contents, whereas the sandstones have high k and low th contents. no distinct 6464 relationship between high u content and mudstone grainsize is seen. ruffell et al. (2003) presented a model for the flux of k, u and th in different weathering systems at basin scale. one of their scenarios is a basin with a low-relief hinterland and a humid climate, which applies to the depositional setting of the atane formation. the model predicts that chemical weathering dominates and that k and u are removed in solution to sea water, while th is concentrated in detrital clay. this model may explain the relatively high th radiation in the detrital mudstones of the atane formation. k-feldspar is a minor constituent of the sandstones but contributes significantly to the k-radiation in the sandstones. the model further predicts that k and u are enriched in authigenic minerals in the basin. such enrichment of u is not observed in the atane formation, possibly due to the relatively high sedimentation rate. the low u content may also reflect the predominance of land plants (type iii kerogen) that generally contain small amounts of u in organic complexes compared to marine organic material. a comparison of sedimentological logs from delta-front successions with their total gamma radiation in a case study from ireland also indicates that the delta-front successions are difficult to identify from the gamma log alone (davies & elliott 1996). summary the present study demonstrates the importance of calibrating petrophysical logs to core data. the cyclicity which characterises the atane formation in outcrops and cores (figs 1, 2), and which would be a means of identifying the atane formation in an un-cored well, is obscured in the total gammaray log. this fails to resolve the grain-size variation in the atane formation because the presence of k-poor kaolinite, despite enrichment by th, provides a ‘cleaner’ signature of the mudstones while the presence of sand-sized k-feldspar gives a ‘dirtier’ signature of the sandstones. the th log can, to some degree, resolve the lithological variation but the energy is too small to be reflected in the total gamma-ray signal. the high sedimentation rate and the predominance of terrestrial organic material precluded the development of a characteristic u signature in the marine mudstones. the gamma log thus shows the variations in elements which occur in small amounts, because the bulk of the sediment (quartz, kaolinite and terrestrial coal debris) contributes very little to the gamma-ray radiation. references chalmers, j.a., pulvertaft, t.c.r., marcussen, c. & pedersen, a.k. 1999: new insight into the structure of the nuussuaq basin, central west greenland. marine and petroleum geology 16, 197–224. dam, g., pedersen, g.k., sønderholm, m., midtgaard, h.m., larsen, l.m., nøhr-hansen, h. & pedersen, a.k. 2009: lithostratigraphy of the cretaceous–paleocene nuussuaq group, nuussuaq basin, west greenland. geological survey of denmark and greenland bulletin 19, 171 p. davies, s.j. & elliott, t. 1996: spectral gamma ray characterisation of high resolution sequence stratigraphy: examples from upper carboniferous fluvio-deltaic systems, county clare, ireland. in: howell. j.a. & aitken, j.f. (eds): high resolution sequence stratigraphy: innovations and applications. geological society (london), special publications 104, 25–35. pedersen, g.k., andersen, l.a., lundsteen, e.b., petersen, h.i., bojesenkoefoed, j.a. & nytoft, h.p. 2006: depositional environments, organic maturity and petroleum potential of the cretaceous coal-bearing atane formation at qullissat, nuussuaq basin, west greenland. journal of petroleum geology 29, 3–26. rider, m.h. 1990: gamma-ray log shape used as a facies indicator: critical analysis of an oversimplified methodology. in: hurst, a., lovell, m.a. & morton, a.c. (eds): geological applications of wireline logs. geological society special publications (london) 48, 27–37. ruffell, a.h., worden, r.h. & evans, r. 2003: palaeoclimate controls on spectral gamma-ray radiation from sandstones. international association of sedimentologists special publication 34, 93–108. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: gkp@geus.dk fig. 3. spectral gamma logs from 40 m of the core from borehole 247801. note the difficulty in interpreting the total gamma radiation log, the negative correlation between k and th content, and the absence of maxima on the u log. the core section is located in fig. 2. yellow: sandstone, grey: mudstone, black: coal or plant debris, black stars: pyrite. cps: counts per second. the total gamma radiation (gr) may be compared to the radiation measured in the borehole (fig. 2). 145 140 120 125 110 115 130 135 105 0 1 2 3 4 0 0 5 10 151 2 3 4 gr total (cps) k (%) 0 5 10 15 20 25 th (ppm)u (ppm) ★ ★ vf clay silt sand f m c mailto:tl@geus.dk geological survey of denmark and greenland bulletin 28, 2013, 33-36 33 assessing urban groundwater table response to climate change and increased stormwater infiltration mark t. randall, lars troldborg, jens christian refsgaard and jacob b. kidmose the global climate is expected to show continued warming throughout the coming century. as a direct consequence of higher temperatures, the hydrological cycle will undergo significant changes in the spatial and temporal distribution of precipitation and evapotranspiration. in addition to more frequent and severe droughts and floods, climate change can affect groundwater recharge rates and groundwater table elevation (bates et al. 2008). some previous studies of climate change impact on groundwater have suggested alarming reductions in groundwater recharge and lowering of water tables. other studies, especially those focusing on regions of higher latitudes, have indicated a potential rise in water tables due to increased precipitation and recharge (scibek & allen 2006; woldeamlak et al. 2007). in addition to changes in precipitation patterns, a shift in stormwater infrastructure design may also alter the hydrologic cycle of urban areas. in recent years, there has been a growing trend towards adoption of low-impact development practices managing stormwater runoff. these practices aim to mitigate the impacts of urbanisation such as increased runoff volume, higher peak runoff flows, lowered water tables and reduced water quality (prince george’s county 1999). in contrast to conventional stormwater infrastructure, which is designed to rapidly collect and convey runoff, low-impact development practices are designed to slow runoff, remove pollutants and evapotranspirate and infiltrate runoff locally. in recent years, numerous modelling studies have investigated the potential impact of stormwater infiltration on groundwater levels. gobel et al. (2004) used a combination of models (gwneu, hydrus-2d, spring) to demonstrate that the installation of infiltration practices across an urban catchment area in germany could raise the groundwater surface by up to 2.3 m in some locations. in another catchment scale study, maimone et al. (2011) used the modelling code dynflow to show that the future groundwater table may eventually stabilise up to 1.5 m higher than its current level in parts of philadelphia, if the city’s plan to alter 40% of its impervious areas into so-called ‘green’ stormwater recharge areas is completed. thompson (2010) used hydrus-2d to demonstrate that a stormwater infiltration basin could cause up to 1.3 m of localised groundwater mounding. in yet another study, endreny & collins (2009) used modflow to show that rain gardens installed throughout a residential catchment area could raise the steady-state groundwater table by up to 1.1 m. the studies mentioned above have investigated groundwater level response to either changes in climate or stormwater management infrastructure. however, to the authors’ knowledge no studies have investigated the concurrent effects of both alterations on the urban hydrologic cycle. in urban areas, it is necessary to determine the potential magnitude of the combined impact, as a steep rise in groundwater level can damage building foundations and subsurface infrastructure due to flooding and buoyancy forces (gobel et al. 2004; vázquez-suñé et al. 2005). this study aims to assess the potential changes in groundwater response caused by both increased precipitation and widespread instalment of stormwater infiltration infrastructure in the city of silkeborg, denmark, using the mike she model. change of groundwater level at the planned location of a new motorway in silkeborg is the focus of this study as portions of the config. 1. the silkeborg study area and the proposed course of the motorway. inset: the location of silkeborg in jylland. © 2013 geus. geological survey of denmark and greenland bulletin 28, 33–36. open access: www.geus.dk/publications/bull 1 kmend of pipe recharge area pervious area impervious area motorway jylland silkeborg c. 9°40´e c. 56°11´n 3434 struction are expected to come critically close to the present high groundwater table in that area. knowledge of the magnitude of potential groundwater changes is essential because improved drainage measures and increased use of concrete will significantly raise the costs of the new motorway. study area the city of silkeborg has a population of c. 43 000 inhabitants and is located in the central part of jylland, denmark (fig. 1). the focus of this study is just north of the river gudenåen, where a portion of the new motorway will be constructed c. 6 m below the present terrain surface. the surficial geology is dominated by coarse-grained, postglacial, sandy sediments that form an upper unconfined aquifer with a vertical extent of 10–15 m. the average precipitation in silkeborg during the period 1961–1990 was 903 mm per year, and the average potential evapotranspiration was 546 mm per year. the average monthly temperature during that period was 15.2°c in july/august and −0.3°c in january/ february (kidmose et al. 2013). methods hydrological models – mike she is a deterministic, fullydistributed and physically based model software capable of simulating surface and subsurface hydrological processes. the danish national water resources model (dk-model) is based on mike she and incorporates national data on geology, soil type, land use, topography, river network geometry, water abstraction and climate. the silkeborg model is a 100 m grid local model using hydraulic head boundary conditions from the 500 m grid dk-model. a 9.2 km2 area within the 103 km2 silkeborg model, which encompasses the new motorway construction and the greater part of the urbanised surroundings, was chosen for the current study (fig. 1). details on the development, calibration and validation of the dk-model and the silkeborg model are found in højberg et al. (2013) and kidmose et al. (2013), respectively. six different model scenarios have been evaluated (table 1). stormwater infiltration modelling – the silkeborg study area consists of 65.5% pervious and 34.5% impervious cover. in the scenarios with conventional drainage stormwater infrastructure (i.e. the ‘cd’ scenarios), 100% of the precipitation on impervious cells was routed directly to the river system (fig. 2a). precipitation on impervious cells had one time step (i.e. one day) to infiltrate or evapotranspirate. at the end of the time step, any water in excess of a detention storage of 4.7 mm (based on calibration results) was routed overland to adjacent cells based on topography. it is assumed that the cd-2010 scenario is representative of silkeborg’s current climate and stormwater conditions. in the end of pipe recharge (epr) scenarios (fig. 2b), it was assumed that 10.7% (34 ha) of the city’s pervious area has been turned into end of pipe stormwater infiltration ponds (figs 1, 2). model cells which were assumed to contain infig. 2. three model scenarios for stormwater drainage infrastructure. scenario name climate data input stormwater infrastructure cd-2010 recorded 1991–2010 conventional drainage to river system epr-2010 recorded 1991–2010 end of pipe infiltration ponds lar-2010 recorded 1991–2010 local area recharge cd-2100 projected 2081–2100 conventional drainage to river system epr-2100 projected 2081–2100 end of pipe infiltration ponds lar-2100 projected 2081–2100 local area recharge table 1. summary of model scenarios a b c conventional drainage end of pipe recharge local area recharge 35 filtration ponds were assigned detention storage of 500 mm to represent the storage depth of the pond. in the epr scenarios, precipitation which would normally be applied to impervious cells was reduced to zero, and the equivalent volume of precipitation was instead evenly distributed over the infiltration pond cells via an increase in precipitation applied to those cells. in the epr scenarios there were 9.3 times as much impervious drainage area as infiltration pond area, so the infiltration pond model cells had 1030% (i.e. 100% + 9.3 × 100%) of the actual rainfall applied to them. this method of manipulating precipitation to simulate the collection of stormwater in specialised infiltration areas on a city-wide scale is similar to the modelling strategy used by holmandodds et al. (2003). the local area recharge (lar) scenarios represent a system where stormwater is managed at the level of individual plots through any combination of infiltration practices, each no more than tens of metres across. it was assumed that infiltration possibilities are numerous and located in close proximity so that at the scale of the model, each cell effectively behaves as a pervious cell. therefore, all paved areas were given properties identical to the pervious areas with infiltration rates controlled by the underlying soils. climate input – precipitation, temperature and evapotranspiration data from the danish meteorological institute from 1991 to 2010 were used as input to the ‘2010’ model scenarios. the input climate data for the ‘2100’ scenarios were generated by applying correction factors based on nine climate model projections from the ensembles project (christensen et al. 2009) to present-day climate data. further information on the delta change downscaling method used can be found in seaby et al. (2013). to generate the results for each of the three ‘2100’ infrastructure scenarios, the model was run nine times (once for each of the nine climate model projections), and the results averaged. results water table elevation – average groundwater elevations along the area planned for the motorway construction were extracted from the mike she model results (fig. 3). areas where the solid black line (i.e. the motorway surface) drops below the water table indicate portions of the motorway which could be flooded by groundwater. in the cd-2010 scenario, a stretch of 160 m of motorway is below the average water table. in the cd-2100 scenario, the average groundwater table elevation is raised by 0.08 m, and the length of motorway surface at risk is extended to 180 m. hundreds of metres of the proposed motorway are potentially flooded in the lar-2010 and the lar-2100 scenarios where the average water table rose 0.48 and 0.55 m above cd-2010 levels, respectively. the highest average water tables of 1.15 and 1.19 m above cd-2010 occur in the epr-2010 and epr-2100 scenarios, which would both put a stretch of nearly 1 km of the proposed motorway at risk. the results indicate that the impact of climate change (i.e. the difference between the ‘2010’ and the ‘2100’ scenarios) is small compared to the impact of extensive implementation of either local area or end of pipe stormwater infiltration practices. only average water tables are presented here to compare the relative impacts of different model scenarios. however, maximum water tables could put much longer sections of the motorway at risk and will therefore be considered in the final design of the motorway. water balance – average yearly volumes of precipitation, evapotranspiration, recharge and overland flow were calculated for the 1991–2010 time period for each stormwater infig. 3. average modelled groundwater table elevations along the 2000 m of projected motorway at silkeborg. the results are relative to cd-2010. model scenario mean (mm/year, 1991–2010) precipitation evapotranspiration recharge overland flow baseflow cd 911 319 304 292 8 lar 911 441 463 11 15 epr 911 311 588 19 29 table 2. catchment water balances for different stormwater infrastructure scenarios lar-2010 lar-2100 epr-2010 epr-2100 motorway surface cd-2100 nw se –0.5 0.5 1.0 1.5 3.0 2.5 2.0 0.0g ro un dw at er ta bl e (m ) 0 500 1000 1500 2000 distance along subsurface motorway stretch (m) 3636 frastructure scenario using mike she’s water balance tool (see table 2). evapotranspiration was greater in the lar scenario, due to the much larger evaporation surface available. recharge was much higher in both infiltration scenarios than in the cd scenario. overland flow, or the volume of water which flows directly into the river system, was very small in both the infiltration scenarios in comparison to the cd scenario which routed all water from impervious areas into the nearest stream. baseflow was highest in the epr scenario, followed by the lar scenario and finally the cd scenario, as would be expected based on the relative recharge volumes in these scenarios. summary and conclusions previous studies have reported groundwater level rise due to either climate change (scibek & allen 2006; woldeamlak et al. 2007) or stormwater infiltration practices (gobel et al. 2004; maimone et al. 2011). however, these two changes to the urban hydrologic cycle are typically not assessed in an integrated way as in this study. the modelling results presented in this paper are within the ranges of the above studies, i.e. tens of centimetres due to climate change and potentially more than 1 m due to the widespread adoption of stormwater infiltration practices. however, these results are specific to the silkeborg motorway and it is expected that the relative magnitude of the impact due to climate change and stormwater infiltration could vary greatly under different climatic and geological regimes. stormwater infiltration practices are often regarded as a form of climate change adaptation in the field of stormwater management as they can help to accommodate the higher intensity and larger volume precipitation events expected in the future. however, as the results of this study indicate, these same practices amplify other problems associated with climate change (i.e. groundwater table rise). the study clearly shows the need for integrated research of urban hydrology, and communication between hydrogeologists, stormwater engineers, planners and policy makers. acknowledgement we thank the danish road directorate for funding this study. references bates, b., kundzewicz, z., wu, s. & palutikof, j. 2008: climate change and water. intergovernmental panel on climate change, technical paper 6, 200 pp. geneva: ipcc. christensen, j.h., rummukainen, m. & lenderink, g. 2009: formulation of very-high-resolution regional climate model ensembles for europe [research theme 3]. ensembles: climate change and its impacts: summary of research and results from the ensembles project, 47–58. exeter, uk: meteorological office hadley centre. endreny, t. & collins, v. 2009: implications of bioretention basin spatial arrangements on stormwater recharge and groundwater mounding. ecological engineering 35, 670–677. gobel, p. et al. 2004: near-natural stormwater management and its effects on the water budget and groundwater surface in urban areas taking account of the hydrogeological conditions. journal of hydrology 299, 267–283. højberg, a.l., troldborg, l., stiesen, s., christensen, b.b.s. & henriksen h.j. 2013: stakeholder driven update and improvement of a national water resources model. environmental modelling & software 40, 202–213. holman-dodds, j.k., bradley, a.a. & potter, k.w. 2003: evaluation of hydrologic benefits of infiltration based urban storm water management. journal of the american water resources association 39, 205–215. kidmose, j., refsgaard, j.c., troldborg, l., seaby, l.p. & escrivà, m.m. 2013: climate change impact on groundwater levels: ensemble modelling of extreme values. hydrology and earth system sciences 17, 1619–1634. maimone, m., o’rourke, d.e., knighton, j.o. & thomas, c.p. 2011: potential impacts of extensive stormwater infiltration in philadelphia. environmental engineer 14, 29–39. prince george’s county 1999: low-impact development design strategies: an integrated design approach, 150 pp. prince george’s county, md: department of environmental resources. http://water.epa.gov/polwaste/green/upload/lidnatl.pdf scibek, j. & allen, d. 2006: comparing modelled responses of two highpermeability, unconfined aquifers to predicted climate change. global and planetary change 50, 50–62. seaby, l.p., refsgaard, j.c., sonnenborg, t.o., stisen, s., christensen, j.h. & jensen, k.h. 2013: assessment of robustness and significance of climate change signals for an ensemble of distribution-based scaled climate projections. journal of hydrology, http://dx.doi.org/10.1016/j. jhydrol.2013.02.015 thompson, a., nimmer, m. & misra, d. 2010: effects of variations in hydrogeological parameters on water-table mounding in sandy loam and loamy sand soils beneath stormwater infiltration basins. hydrogeology journal 18, 501–508. vázquez-suñé, e., sanchez-vila, x. & carrera, j. 2005: introductory review of specific factors influencing urban groundwater, an emerging branch of hydrogeology, with reference to barcelona, spain. hydrogeology journal 13, 522–533. woldeamlak, s., batelaan, o. & de smedt, f. 2007: effects of climate change on the groundwater system in the grote-nete catchment, belgium. hydrogeology journal 15, 891–901. authors’ addresses m.t.r., computational hydraulics international, 147 wyndham street north, suite 202, guelph, ontario, n1h 4e9 canada. e-mail: mark@chiwater.com l.t., j.c.r. & j.b.k., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. http://dx.doi.org/10.1016/j.jhydrol.2013.02.015 http://dx.doi.org/10.1016/j.jhydrol.2013.02.015 geological survey of denmark and greenland bulletin 3, 98-112 98 the skagen well the skagen well – perspectives the perspectives of the drilling of the skagen well can be seen by describing the sedimental changes observed in the succession of strata penetrated. however, for the first 30 m of skagen well iii, only wash-samples were taken, in order to establish the well for further drilling down to the prequaternary. therefore, the first drilling segment, composed of sand and gravel, was later repeated elsewhere in order to obtain core samples also from the topmost part. this was done in a nearby position – the so-called skagen iv well – and consequently the full record of shell material and sediments can be given from the core samples obtained from the well, representing all the strata met with in the skagen area from the quaternary, the skagen iii well dgu file no. 1.287. the skagen wells iii and iv are considered to represent one well and are therefore listed together. however, also the wash-samples obtained throughout the quaternary are represented, but only as qualitative analyses with the first occurrences of macrofossils – especially the molluscs – indicated (appendix 2). thirty metres below surface (+ 1 m a.s.), the sediment is well-sorted fine sand. for the following 10 m to 40 m b.s., the average grain size falls within the coarse silt fraction which is consistent down to 75 m b.s. as shown from 11 interjacent measurements, appendix 3. from the depth of 75 m b.s. the average diameter falls within that of medium silt, down to the level of 100 m b.s., still well sorted. at the following levels to a depth of 133 m b.s., the material is fine silt and clay. at a depth of 135 m b.s., fine sand with poor sorting occurs, and at 136–137 m b.s., with very poor sorting and quartiles 40% / 90% of 54.099, which shows a diamict material with a content of stones and only allochthonous shell material, in contrast to the superjacent 130 m. at a depth of 179 m, the well produced a fine-grained material of medium silt, moderately sorted, which is close to what was found above the diamict sequence. with some rise in the average diameter to fine sand and a lowering of the sorting, the next remarkable shift happens at a depth of 187 m b.s., where a new diamict sequence is found down to 195 m b.s. here the boundary to the prequaternary deposits is met with, most probably belonging to the lower cretaceous, according to skagen ii, dgu file no. 1, 43 (sorgenfrei & buch 1964). from the above-mentioned strata in the cored sections, the well can be divided into two parts from the point of view of the present investigation on macrofossils. a sandy to gravelly, clayey to silty wellsorted material found in the upper 130 m of the well and in a smaller interval of almost 10 m at a depth of 179 to 187 m b.s. between these two parts, diamict and well-sorted clayey layers without an in situ macrofauna are found. as the main characteristic feature of the 140 m finegrained and well-sorted material, its content of shellbearing marine molluscs is considered. however, also other marine macrofossils have been recognised but not referred to species level, although recorded on a higher taxonomic level in appendix 3, which represents the finds in the skagen well. also the sedimental data are all shown in appendix 3, allowing a direct comparison between the finds of faunal elements and lithology sensu lato. the organic compound is shown with loss on ignition (550°c) and the occurrences of concretions such as pyrite and iron compounds. also allochthonous shell materials are figured. in consequence of the quantitative analyses of molluscs, diversity and number of specimens are given for the marine strata throughout the whole sequence. it is shown that these figures are very greatly according to the different facies met with. on the basis of the dating within the limits of the carbon-14 method (heier-nielsen et al. 1995), appendix 4 and fig. 3, it is seen that the 130 m thick upper sequence of the well is dated to the last 15 000 14c years. this comprises the whole of the holocene with its boreal sea deposits and the late weichselian with the arctic younger yoldia deposits. but also from 140–150 m b.s., measurements on gas compounds of marine origin (t. layer, personal communication 1999) and shell fragments have been dated, giving ages around 36 000 b.p. 14c years. this means that the diamictic sequence occurring between the two marine strata has taken up material which in age is equivalent to the younger part of the older yoldia clay sediments. the deposition of the diamictic geus bulletin no 3.pmd 28-06-2004, 08:4598 99 sediments is referred to the time of the late weichselian ice advance to the main stationary line in jylland. molluscan shells (macoma calcarea) from the younger part of the older yoldia clay (the macoma calcarea zone, sensu petersen in bahnson et al. (1974)), has recently been dated by the ams method to be in 14c years around 32 000 – 33 000 b.p. (aar-1410 and aar-1411). consequently, it is likely that the older marine strata in the lower marine part of the skagen well can be correlated to part of the sequence demonstrated in the skærumhede well (jessen et al. 1910; bahnson et al. 1974) covering the eemian and the main part of the weichselian. regarding the 130 m thick upper sequence, this is from results of the 14c dates related to the time of the younger yoldia sea and the holocene, and as it appears from the dates of the sediments here is for the first time within the danish area found in a continuous marine succession. this can be explained on the basis of the hitherto unsurpassed thickness of late weichselian–holocene marine strata. therefore, while most of the danish area has a continental period in the time span 11 000 – 7500 b.p. in 14c years (petersen 1985b), the skagen area was so low-lying that it was continuously covered by the sea. this is a reflection of a lesser glacial deposition and the glacio-isostatic down-pressing – the latter amounting to up to 200 m, as seen from the amount of isostatic rebound after the waning of the ice cap over northern denmark (petersen 1990). however, when only holocene dates are used, the estimated rebound of 200 m seems to be greater than the present dates allow when also late weichselian dates are used (petersen 1999). so the low stand of sea level during the latest part of theweichselian and early holocene was not to be overtaken by the isostatic uplift at any time in this area. after the final large eustatic rise in the early atlantic, the marine sedimentation is a dominating factor in raising the levelled sea bottom, compared to a decreasing isostatic rebound up to the subatlantic when the isostatic rebound in denmark expired (petersen 1990). here the formation of the skagen spit takes over, so that the last event of changing depth depends on the large quantities of sand and gravel deposited as the skagen spit grows to the north. the pre-late quaternary deposits in the deeper part of the skagen well, the base of the quaternary is found resting on pre-quaternary deposits of lower cretaceous sand. in the following description, appendix 3 should be consulted. the pre-quaternary strata consist of quartz sand and gravel. from a depth of 195.15–195.30 m b.s., which is the upper part of the pre-quaternary stratum, a mean grain size of fine sand, poorly sorted, is found (lab. no. 789.93). at a depth of 194.35–194.48 m b.s., the sediment is poorly sorted and the mean grain size is within medium sand, and this sediment also contains much quartz, but has another component in the form of stones of granitic composition. the cumulative curve shows an even distribution of all grain sizes, which refers this sediment to be a till (lab. no. 788.93). this is also true for the overlying strata up to around 189 m b.s. the mean grain size is here within fine sand; however, more fine-grained parts are found. the sediment contains siderite(?) concretions with pyrite, in which imprints of cyrena sp. are found. this might show that jurassic deposits have been eroded. some traces in pyrite were found as well. the whole sequence can be regarded as till. the sediment analyses from 188.57 to 187.18 give a badly sorted sediment sustaining that this is a till. also this level contains shell fragments, one of which can be shown to be a nuculana pernula. the late pleistocene eemian deposits the granulometric composition is shown from sample lab. no. 800.93. at a depth of 185.37 m b.s. the sediment is very fine-grained but contains only fragmented bivalves. at 185.0 m b.s. the marine sediment can be demonstrated by the occurrences of dentalium vulgare in many specimens and the bivalve kelliella miliaris also in many specimens and with connected valves. the granulometric composition can be seen from the two levels 182.65 and 180.57 m b.s., samples nos 784.93 and 797.93 respectively. it appears that the sediment is very fine-grained clay to fine silt and moderately sorted in the 180.57 m level. accessory finds of spatangids and ophiorids occur at the 185.0 m level, and pyrite formed in former burrows in the clayey material. also finds of fish occur, as found at the 182 m level, geus bulletin no 3.pmd 28-06-2004, 08:4599 100 and under the name of other fossils also crustacean remains have been listed. the third mollusc species found at this level is the ophistobranch limacina retroversa, which is found in large numbers (11 specimens in one sample) together with kelliella miliaris (also of a number of 15 in one sample). the samples here referred to are all from the heavy weight separation of the foraminifer samples. the species diversity and number of specimens in the sediment appear from the sample at 182.24 m b.s. in which 25 specimens of kelliella miliaris are found – most of them with connected valves and in some parts kept in pyrite. trace fossils in pyrite are found in great quantities recorded in the table from all levels. an expression of the grade of fine-grained sediment occurring at this level can be seen from the fact that only a biogene residue occurs here including the pyritiferous biogene traces – lebenspuren. the limpid delectopecten vitreus also appears at this level. as mentioned in the chapter on the molluscan species, the two sedentary species which today are known from the deeper part of the skagerrak are delectopecten vitreus and kelliella miliaris. the latter is also found in the turritella terebra zone in the skærumhede well. from 183.4 m b.s. hiatella arctica is found, which occurs also at the greater depths and furthermore is a species widely extended. the occurrence of entalina tetragonia at 183.6 m b.s. goes together with the occurrence of delectopecten vitreus, both of which are found in the deeper part of the skagerrak today, where they are part of the amphilopsis norwegica/delectopecten vitreus community. to this can be added cadulus jeffreysi found at the 184.4 m level. this species is widely extended in the northern part of the atlantic down to the bay of biscay and into the mediterranean. however, a single find of siphonodentalium lobatum at 184.6 m b.s. points to a more arctic environment. such shells are found in glaciation deposits according to muus (1959). however, the species may extend into the lusitanian region. from the 182 m level and up to 180 m b.s. still with a fine-grained and well-sorted sediment, yoldiella frigida appears, which is also known from the deeper part of the skagerrak today. this species can be referred to the same environment as has been mentioned above – the amphilopsis norwegica/delectopecten vitreus community. yoldiella frigida is known from the turritella terebra zone in the skærumhede sequence and the portlandia arctica zone according to jensen & spärck (1934). kelliella miliaris and limacina retroversa, which have been very frequent in marine layers met with under 180 m b.s., are no longer found above 180 m b.s. the early/middle weichselian, marine and glacigene deposits regarding the sediment which is to follow at the levels above, between 179.65 and 179.74, it appears that the mean grain size is somewhat bigger medium sand, moderately sorted. but the most significant is found in the cumulative curve showing two maxima on the frequency curve (fig. 91). this points to the effects of two sedimentation agents which might be a drop till effect besides the general marine sedimentation. during the examination of the samples from this level, sand and fine gravel occur, in contrast to the levels below, where only biogene remains were found, including the pyritiferous biogeneous traces. the coarser minerogene elements are found higher up in the series to the level 175.30–175.50 m where a fine-grained sediment with a median grain size of 0.002 mm reveals arctic marine molluscs. this is the first appearance of portlandia arctica, which as the name tells is the characteristic mollusc of the portlandia arctica zone in the skærumhede sequence. however, also the presence of yoldia hyperborea, which is known today from the arctic and down to the lofoten area is characteristic. this species is also found in the portlandia arctica zone together with nuculana pernula and palliolum greenlandicum. the occurrence of macoma sp. has been added from cumulated weight per cent frequency per cent grain size distribution older yoldia sea sediment sample id: 179.65 – 179.74 m 100 90 80 70 60 50 40 30 20 10 0 0. 00 20 0. 00 28 0. 00 39 0. 00 55 0. 00 78 0. 01 10 0. 01 56 0. 02 21 0. 03 12 0. 04 42 0. 09 00 0. 12 50 0. 18 00 0. 25 00 0. 35 50 0. 50 00 0. 71 00 1. 00 00 0. 06 30 w ei gh t pe r ce nt , % grain size, mm fig. 91. the granulometric composition with two maxima on the frequency curve (lab. no. 483.93) at the 179.65–179.74m level. geus bulletin no 3.pmd 28-06-2004, 08:45100 101 the 177.8 m level but not on species level because of the fragmentary state of the shell. both spatangids and ophiorids are found and a fragment of cirriped at the 177.8 m level. at 174.4 m b.s. a single find of yoldia hyperborea occurs. the granulometric composition found at the 173.67–173.85 level (lab. no. 781.93) shows bad sorting in a clayey sediment with a median grain size within clay to fine silt. in this sediment fragments of arctica islandica are found that can be regarded as being part of the redeposited material which can be found also higher up in the core. at a level of 166.5 m b.s. the sediment is well-sorted fine sand and regarded as fluvial. in the following 16 m up to 151.50 m b.s. the sediment is coarser, being a moderately sorted medium sand also regarded as fluvial sand. from here only some shell fragments are found and no record of fossilia varia (other fossils in appendix 3). in the next metres to the level of 143.83–144.00 m b.s. the mean grain size is within the medium silt grade. this is found to be a fine-grained fluvial material forming part of a glacigene complex. also here, fragmentary molluscs are found. the late weichselian marine and glacigene deposits the first molluscs regarded as autochthonous above the glacigene complex are found at the 130.2 m level, so this is regarded as the upper limit of the glacigene sequence. in the interval from 141.00 and up to 130.2 m b.s. more shell fragments have been found – all showing signs of transport and wear. finds of pyrite (137.8 m level), concretion (132.6 m level), and glacial stria on a stone (137.44 m level) all reflect typical features for a till deposit. the marine shell material taken up by the glacier occurs in a fragmentary state, which is typical for redeposited material. however, it is from these strata that the absorption of gases from marine deposits has been dated. these dates form as mentioned a parallel to the age determination of the shells (macoma calcarea) from the skærumhede sequence where the older yoldia clay fauna has been studied (madsen et al. 1908; bahnson et al. 1974). the ages found on macoma calcarea shells from the skærumhede ii well give for the first time, on the basis of molluscs, the absolute age of around 32 000 – 33 000 (14c years) of the younger part of the older yoldia clay. compared with the dating of the marine gases from the skagen well, there is a good correlation to the stratigraphically now well-established skærumhede sequence, so that the two cored sections found on skagen and at skærumhede can be regarded as deposited during the same time in the weichselian. the skagen sequence, however, has been strongly eroded by the ice sheet advancing during the late weichselian. however, the thick packet of up to 50 m glacial sediments consequently contains the traces (gases) of that marine environment, which has been eroded, but is hereby dated to give the maximum age of the glaciation. this age points to the glaciation event in the late weichselian around 20 000 – 18 000 b.p. (petersen & kronborg 1991). however, here the upper marine sequence found in the skagen well will be described. from the 131.63–131.73 m level and up the core the sediment is extremely fine-grained with a medium grain size of fine silt which stays as such a size up to 100 m b.s. it should be noted that throughout the first 15 m of the core from the above-mentioned level finds of coarser material occur. this is seen at the 125.89–126.00 m level (lab. no. 526.93), where the granulometric composition reflects two maxima on the frequency curve (fig. 92). this is taken as a typical sign of a supplementary sedimentation which might have been caused by the melting of floating ice with the coarser sediment imbedded – a drop till effect, as found deeper in the core (fig. 91). this suspected arctic influence is sustained by the occurrences of arctic molluscs up to the 114.0–115 m level, where both portlandia arctica and bathyarca glacialis are present. cumulated weight per cent frequency per cent grain size distribution younger yoldia sea sediment sample id: 125.89 – 126.00 m 100 90 80 70 60 50 40 30 20 10 0 0. 00 20 0. 00 28 0. 00 39 0. 00 55 0. 00 78 0. 01 10 0. 01 56 0. 02 21 0. 03 12 0. 04 42 0. 09 00 0. 12 50 0. 18 00 0. 25 00 0. 35 50 0. 50 00 0. 71 00 1. 00 00 0. 06 30 w ei gh t pe r ce nt , % grain size, mm fig. 92. the granulometric composition with two maxima on the frequency curve (lab. no. 526.93) at the 125.9–126.0 m level. geus bulletin no 3.pmd 28-06-2004, 08:45101 102 furthermore, species such as nuculana pernula, nuculana minuta and yoldiella lenticula occur, which are known from the older yoldia clay in the skærumhede sequence. yoldiella frigida is the first to occur at the 130.2 m level in the skagen well. from this level several finds of ophiuroids (fragments), cirripeds and remains of pisces. however, no finds of spatangoids have been demonstrated within the whole sequence referred to the arctic marine deposits, but they are found all the way up in the boreal sequence (fig. 93e; fold-out, back cover). in the upper part of the arctic sequence siphonodentalium lobatum occurs at the 116.0–114.6 m level and a single find of entalina tetragona. occurrences of nucula sp. and macoma sp. are also recorded in the arctic part, but in such a fragmented state that the species cannot be given. from the recorded faunal composition it appears that it is a deeper-water fauna. this is also supported by the fact that forms reflecting an arctic macoma calcarea community are not present, and the fine-grained sediment points in the same direction. as a comment to the sedimentary environment it should be mentioned that magnetic spherical concretions have been found all through the arctic sequence. from five levels: 117.69–117.85, 124.34–124.50, 127.39– 127.50, 128.13–128.33 and 132.69–132.77 m b.s. a high content of griegite (fe 3 s 4 ), which explains their magnetic quality, has been found by x-ray analysis together with quartz, calcite, feldspar and clay. griegite has been reported as a constituent of reduced sediments. the occurrences in the skagen well are of interest in so far as the spherical magnetic concretions have been recorded only from the arctic sequence. this arctic sea deposit has been dated on material from the cores both foraminifers and macrofossils (heier-nielsen et al. 1995). from this it is seen that the actual time span ranges over 5000 14c years from 15 000 to 10 000 b.p. the sudden change in the macrofauna, or better the abrupt stop of the occurrences of arctic species, at the level of 114.2 to 114.00 m b.s. gives the pleistocene– holocene boundary. the transition from the pleistocene younger yoldia sea to the holocene marine deposits is here recorded for the first time within the danish realm with a whole series of ams datings supporting the chronostratigraphic position, see appendix 4. the dates are highly significant because the molluscan finds in the older part of the marine holocene are extremely poor. this, however, is not caused by the lack of samples from this core section, but is as will be shown dependent on the type of facies following the deposition of the youngest yoldia sea, which was a deeper-water deposit, followed by a boreal deeperwater facies in the older part of the holocene. the change from arctic to boreal conditions is regarded as influenced by a new current system from the atlantic bringing in the new temperate fauna replacing the arctic fauna of late weichselian age. the change in fauna is, however, not reflected in the sedimentary record (appendix 3, pp. 17, 21), which shows a very homogeneous clayey grain size distribution with nothing coarser than fine sand. only in one sample (appendix 3, p. 21, 115 m b.s.) at the sharp boundary between late weichselian and holocene medium sand, coarse sand and gravel are observed. on this homogeneous sequence of clay to fine sand measurements of magnetic susceptibility and thermoluminescence sensitivity have been conducted. it is worth noticing that in a diagram of magnetic susceptibility versus tl sensitivity the two samples forming the peak in the last part of the late weichselian also represent the more immature sediment (high susceptibility and high tl sensitivity). in contrast, the whole series of samples from the lower part of the holocene seems more mature (low susceptibility and low tl sensitivity). so, in this way the peak can also be connected with the sudden break through of the water from the baltic ice lake at mt. billingen, whereas the mature sediments from the holocene may reflect the long-transported sediments introduced by the new current system from the atlantic, bringing in the new temperate fauna in the early part of the holocene and replacing the arctic fauna of late weichselian age (unpublished data, k.l. rasmussen and k.s. petersen). the holocene as mentioned earlier, the transition from the arctic younger yoldia sea to the oldest holocene marine deposits is not to be seen from the sediment analyses except for the occurrences of griegite and some coarser material in the arctic part. this appears when the cumulative curves from the 125.89–126.0 and 113.60– 113.70 m levels from the arctic and boreal part (figs 92, 94, lab. nos 526.93 and 522.93 respectively) are compared. the median grain size is for both samples fine silt, see appendix 3, p. 21. considering the many samples analysed within the geus bulletin no 3.pmd 28-06-2004, 08:45102 103 lower part of the holocene up to the 100 m level, which is dated to be around the boreal–atlantic boundary, only very few molluscan species have been found; also the number of specimens is low. the preboreal–boreal 10 000 – 8000 14c years b.p. in the preboreal–boreal sequence only parvicardium minimum has been found in more than a single find together with mysella bidentata. however, three other genera are recorded: cardium, abra and lyonsia. parvicardium minimum is known from the deeper part of the skagerrak today and is found up to a depth of 30 m in the kattegat. it is recorded also from the eemian in the skærumhede series. compared with the occurrences of mysella bidentata also in this core level at skagen one can imagine a deeper-water environment, because mysella bidentata is also found to great depth (600 m) today in the skagerrak. spatangoids, apparently in great quantities – considering the many fragments – are found and in a lesser degree fragments of ophiuroids, which were also recorded from the arctic part. from the family spatangidae, five genera are known in nordic waters. from the skagen well at a depth of 108.34–108.56 m b.s. a well-preserved species of brissopsis lyrifera (forbes) has been collected (fig. 95). this species lives only on pure muddy bottoms and totally embedded in the sediment. as seen from the grain-size distribution from the level of 107.90–108.00 m b.s., this part is a fine-grained sediment. from the 109.39–109.50 m level the core section revealed a cut through the traces of a spatangoid similar to those that brissopsis lyrifera could leave, with the typical backfilling (bromley 1990, fig. 5.11; see fig. 96). brissopsis lyrifera can be found in great quantities in the northern part of the kattegat and skagerrak, while it might be found in the øresund but not in the bælt sea, the baltic and the limfjord region, according to mortensen (1924). cumulated weight per cent frequency per cent grain size distribution early holocene sediment sample id: 113.60 – 113.70 m 100 90 80 70 60 50 40 30 20 10 0 0. 00 20 0. 00 28 0. 00 39 0. 00 55 0. 00 78 0. 01 10 0. 01 56 0. 02 21 0. 03 12 0. 04 42 0. 09 00 0. 12 50 0. 18 00 0. 25 00 0. 35 50 0. 50 00 0. 71 00 1. 00 00 0. 06 30 w ei gh t pe r ce nt , % grain size, mm fig. 94. the cumulative curve from the 113.60–113.70 m level, lab. no. 522.93. fig. 95. the well-preserved brissopsis lyrifera (forbes) from the 108.34–108.56 m level. mguh 25404. fig. 96. trace from the 109.39–109.50 m level. might be similar to that of brissopsis lyrifera. geus bulletin no 3.pmd 28-06-2004, 08:45103 104 also other spatangoids might be found in the skagen cores in the huge material of fragments. from the older strata the genus echinocardium has been recorded earlier from the skærumhede series by the author, and echinocardium cordatum has been found in the cyprina clay from the eemian (madsen et al. 1908). so poor in molluscan species this community from the early holocene appears to be, one may pay attention to the abundant of remains of starfishes and echinoids which can be seen as a dominating element in this environment. in this way the sea bottom of those days was controlled by the echinoderms eating up most of the larvae of molluscs, as described by thorson (1961). if one should be compared with a present-day community, it must be the maldane-ophiura sarsi community in which besides ophiura sarsi, brissopsis lyrifera is found as the only often found larger animal (thorson 1968). the maldane-ophiura sarsi community replaces the amphiura community at depths of around 150 m and deeper in the skagerrak. a single find of pisces (100.3–100.5 m) has been recorded and a few finds of plant remains and pyritified traces (chondrites?). these rare finds of marine deeper-water facies from the very last part of the pleistocene and the earliest holocene will contribute to our knowledge of the land and sea configuration during the so-called continental period (petersen 1985b). considering the sedimentation rate during the first 2000 years of the holocene, viz.: through the preboreal and the boreal from which there have been only a few records earlier within the danish area, it is seen to be around 7.5 m per 1000 years. this is higher than the sedimentation rate for the younger yoldia sea, as found also in the skagen well record of 3 m per 1000 years. when this is given in calendar years, the differences are even bigger, because then the sedimentation of 15 m in the younger yoldia sea took about 6000 calendar years and still about 2000 calendar years in the preboreal and boreal seas within the skagen area (petersen & rasmussen 1995a, b). regarding the sediment, 50% is found to be clay in the younger yoldia sea – and in some parts at the level of 117.29–117.40 m b.s. around 63% – while during the preboreal–boreal the clay content has fallen from around 40% at the 113.60–113.70 m level to 20– 30% at the 100 m level. the atlantic 8000–5000 14c years b.p. from the dates (heier-nielsen et al. 1995) the atlantic covers the cored section from 100 to 80 m b.s. here the sediment in the oldest part has 30–20% clay, falling to a content of 15% clay in the youngest part at the 80.60–80.70 m level (see fig. 93a). throughout the atlantic the echinoids still dominate the samples and among these the spartangoids, as in the preboreal and boreal. however, here small gastropodsoccur:melanella lubrica,odostomia umbilicaris, and eulimella scillae. melanella lubrica is regarded as an ectoparasite on holothuroids, and odostomia umbilicaris is often found together with mytilus adriaticus. however, the latter bivalve has not been found in the skagen well material. it should be mentioned that the odostomia species are difficult to determine (fretter et al. 1986, p. 605) and no less so in subfossil material. furthermore,onoba vitrea and aclis minor are found in the younger part of the atlantic, where the determination of onoba vitrea is taken with some reservation because the three other species odostomia semicostata, odostomia aculeus, and odostomia proxima are very much alike and difficult to tell apart on shell features alone. aclis minor belongs to a large group of predatory gastropods that mostly and perhaps always (cf. fretter & graham 1962) are associated with echinoderms. from the atlantic single finds of parvicardium minimum from the 96 m level and spisula subtruncata at the 86 m level occur. spisula subtruncata is found next at the 73 m level in the subboreal, but becomes the dominating bivalve at the 30 m level, which can be referred to the younger part of the subatlantic. this depth is also within the range where this bivalve is found in large amount in the present-day danish seas. from the atlantic the predatory gastropod lunatia alderi occurs. this species is most probably the one which has bored into the many molluscs found in the overlying strata, but has not been recognised by its traces in the sparse material from the atlantic. fragments of abra sp. and macoma sp. occur in the cored section from the atlantic, and from the 80.60– 80.70 m level also finds of pisces and crustaceans have been recorded, as seen in appendix 3. it appears that also in the atlantic the sampling reveals a deposit with low diversity and few specimens of molluscs, where the echinoderms dominate as in the preboreal and the boreal sequence. however, considering the older holocene deposits which were tengeus bulletin no 3.pmd 28-06-2004, 08:45104 105 tatively referred to the maldane-ophiura sarsi community, the one from the atlantic can on the basis of the molluscs and the still dominating echinoderms be regarded as another of the deeper-water communities found in the present-day deeper water in skagerrak. here it should be the amphiura community, which as mentioned earlier is found above depths of 150 m. the subboreal 5000–2500 14c years b.p. the following 20 m of the skagen well cover the subboreal, 80 to 60 m b.s. the sedimentation rate can be estimated to be 8 m per 1000 years, a slight rise from the 6.6 m per 1000 years found during the atlantic. the clay content falls in this part to below 10%, and the coarse silt and fine sand fractions become the dominating grain sizes. thematerial iswell sorted. fig. 93a,c. in all the sampled cores within this section fragments of echinoderms occur – mostly spatangoids as found earlier – but the diversity of mollusc species is higher, up to 10 different species in one sample and several with five to seven species in each. however, the number of specimens is still low and most of the finds are of single specimens. only onoba vitrea is found in a number of eight specimens in one sample (the 67.0 m level). among the other species, only lunatia alderi can be mentioned occurring in a number of eight within the whole section. from the 73.0 m level turritella communis occurs with boring of predatory gastropods – probably lunatia alderi. furthermore, eulimella scillae and retusa truncatula are found from the 67 m level and mysella bidentata together with corbula gibba at the 65.6 m level. the latter will be more common in the above-lying strata belonging to the subatlantic. from the subboreal sequence, one of the very few finds of polyplacophora occurs sitting in the sediment, represented, however, only by one plate which does not allow further determination by the author. within the interval from 75.0 m to 72.0 m three finds of turritella communis have been recorded. this is one of the characteristic species on the level muddy bottoms. nuculana minuta which has been found also in the arctic younger yoldia clay is here recorded for the first time in the holocene in the skagen well. there are several finds of nuculana minuta from the subboreal, and it is found in the present-day kattegat on muddy bottoms at depths of more than 20 m. this fits very well with the occurrences of turritella communis. acanthocardia echinata is also found for the first time and here recorded from 78.0 m. this species occurs on mixed bottoms and clay bottoms at depths of from 10 to 150 m. also phaxas pellucidus occurs for the first time in the skagen well during the subboreal. this mollusc occurs in general at depths of between 10 and 50 m, often together with abra alba, also found in this section of the well. the first occurrences of chamelea striatula and corbula gibba are in the skagen well during the subboreal. chamelea striatula is one of the most common of the danish marine bivalves but is connected to the sandy bottoms. according to jensen & spärck (1934), it is not found in the kattegat at depths greater than 50 m, because the sandy bottom in this region goes no further out and the species is rarely taken on clayey bottoms. in this connection it should be noticed that just around the 75 m level, where chamelea striatula occurs for the first time in the skagen well, the sediment changes to coarse silt with more than 50% fine sand. finally, at 61.09–61.14 m, is the first occurrence of tellimya ferruginosa. this species will also be more common in the subatlantic from the 30 m level. tellimya ferruginosa is often connected with the occurrence of echinocardium cordatum but can also be found free living (jensen & spärck 1934). the many new species – new through time in the skagen well – introduced in the subboreal point to water depths around 50 m with characteristic species from the present-day community such as turritella communis and chamelea striatula – the venus community. the changes to a more sandy sediment are perhaps the background for the occurrences of the new species. however, the echinoids have also decreased – and this may explain the more prolific mollusc faunas, for the toll of eaten molluscan larvae taken by the echinoderms is no longer so high (cf. thorson 1961). the subatlantic 2500– 14c years b.p. the uppermost 60 m of the skagen well belongs to the subatlantic. in general the 60 m cored section that falls within the subatlantic can be divided into two parts of an equal length of 30 m: the lower 30 m with geus bulletin no 3.pmd 28-06-2004, 08:45105 106 a clay content of 30–40% of well-sorted sediment, and the upper 30 m mainly consisting of fine to medium sand with few intercalations of gravel. regarding the dated part of this upper sequence (heier-nielsen et al. 1995, table 1) – from 30.25 to 12.75 m b.s. the sedimentation is 17.50 m during 210 years from a.d. 950 to a.d. 1160. this gives a sedimentation rate of about 80 m per 1000 years. in this case showing a fine example of the building up of the skagen spit, where the coarser material occurs as part of the long-shore transport, and deposited in foreset beds. the older subatlantic the older part of the subatlantic covers the interval from 60 to 30 m b.s. this section shows a slight coarsening upwards and a sedimentation rate of 30 m per 1000 years. the faunal composition can be analysed on the basis of 50 samples with a higher species diversity than found in the subboreal. some of the species are new to the record from the skagen well. hinia pygmaea appears for the first time at the 42 m level, as well as hinia reticulata. they both belong to the sublittoral zone and are found on muddy bottoms. mangelia brachystoma, which first occurred at the 58 m level, also belongs to the sublittoral fauna, but it occurs on sand and sandy muddy bottoms. polygireulima sinuosa is an ectoparasite on echinoderms which are still common and constitute a part of every one of the samples, but it has been found only within the level 38.19–38.24 m. the first littoral species, mytilus edulis, occurs at 49.14–49.24 m, and from this level it occurs regularly upwards, but only in small numbers until the 31 m level, where it is found in greater quantities. this species can be found out to 40 m depth, but must nevertheless be considered an eulittoral species where its occurrence is most abundant. the young specimens are often found on the vegetation. also chlamys varia is common in the coastal zone and occurs at the 32.85–32.90 m level. heteranomia squamula is epifaunal on hard substrates but also on algae and crustaceans. it has a wide occurrence from the littoral zone out to a depth of 100 m. in the skagen well it is confined to the subatlantic part. thyasira flexuosa is today a common bivalve on clayey bottoms from 20 m to 100 m, but it has been found only in two samples from the older subatlantic. this is hard to explain, as it has a wide extension within the whole of the north atlantic area (jensen & spärck 1934), and in numbers it is one of the most dominant species on the muddy bottoms which according to the sediment analysis have been prevalent for most of the holocene in the skagen area. turtonia minuta, belonging to the species from the coastal zone, is found in a single specimen at 39.85– 40.02 m. it is not recorded from the recent danish fauna, but lives off the norwegian west coast and is found subfossil in the limfjord region. at 47.30–47.35 m is the youngest record of parvicardium minimum, which was one of the few species occurring in the older holocene reflecting deeper water. a single find of angulus tenuis is at 55.30–55.35 m. the common occurrence of this species starts at the 30 m level. also donax vittatus occurs at 35.90–36.00 m level which must be seen as outside the general occurrence of this species, which is from littoral to around the 20 m depth donax vittatus is found within the high energy zone. a single find of abra prismatica at 49.90–42.00 m is within the general depth interval for this species (20– 60 m). in connection with the depth indications given in the well in metres below surface and the common depth intervals indicated by various authors for the molluscan species, it is possible to use the actual depth recorded in the well as the living depth for the subfossil fauna found in the skagen well during the younger part of the holocene. this because of the expiring isostatic movement and only little eustatic changes during the late holocene (petersen 1991b). corbula gibba, which was also found during the subboreal has in the subatlantic an even occurrence through the older part. barnea candida, normally only found out to a depth of 30 m, occurs in the well already at the 51.54–51.59 m level, although only found in fragments. cochlodesma praetenue which was found in the eemian at 183.77–184.00 m b.s. is also found in the subatlantic at 43.19–43.24 m. this species is rare in danish waters and has been taken alive only once north-east of the island of læsø. however, shells have been found elsewhere in the kattegat region, jensen & spärck (1934). it has a wide occurrence from the littoral zone and out to 110 m on different bottom types. finally thracia phaseolina shall be mentioned. this species occurs to depths around 50 m on clayey bottoms. as mentioned above, the echinoderms are also found geus bulletin no 3.pmd 28-06-2004, 08:45106 107 in the subatlantic represented by the fragments of spatangoids. also cirripeds occur in still higher quantities up towards the 30 m level (appendix 3, p. 6). furthermore, there are single finds of pisces and other fossil remains such as crustaceans (other fossils in appendix 3). however, also serpulae are found and may have settled on the shells of the other animals as the crustacean carapax. the younger subatlantic the increasing number of cirripeds in the upper 30 m should probably be regarded as allochthonous, since they occur together with the coarser material during the formation of the advancing skagen spit. the change in the upper 30 m to coarser material also introduces new forms of molluscs that are characteristic of the littoral facies and high-energy coastal situation still prevailing in this area today. the description of the upper 30 m is, as mentioned earlier, based on the skagen iv well 50 m away from the skagen iii well and at the same level (+ 1 m). this was done because only washed samples were obtained from the upper 30 m of the skagen iii well, and such samples could not form the best basis for a uniform description of the whole sequence – especially the necessary quantitative treatment of the molluscan faunas could not be fulfilled in that way. furthermore, a total of 29 grain-size analyses have been made within this part of the column, showing two sequences of well-sorted sediment coarsening upwards, with a sorting coefficient lower than 2 (fig. 97). in order to control the degree of transported shell material, size analyses and counts on right and left valves have been considered relevant with such a highenergy near shore sedimentation (fig. 98). especially the most prevalent bivalve within these uppermost 30 m, spisula subtruncata, has been counted. also observed borings have been figured in appendix 3, to be seen in connection with the actual finds of the predatory gastropods. this is done in order to show the degree of mutual connection in the molluscan assemblages, between predatory elements and their prey. the building-up of the upper 30 m took place within a very short period of time, and the sedimentation rate of this interval is estimated to be around 70 m per 1000 years. this high sedimentation rate has a serious effect on animal life. thedates on the building-up of the skagen spit lead to the conclusion that the extension of the coast line to the place where the wells have been sunk took place around a.d. 1400. taking into account that the final history of the coastal development takes place as a near-shore and littoral deposition history, the actual development on a west coast site similar to the skagen area has been analysed. this has been done by way of several van veen grab samples – altogether 61 outside the agger tange complex in the westernmost part of the limfjord (petersen 1994a). these investigations focused on the bivalves, evaluating their degree of being autochthonous from the preservation with both valves together, one valve but whole, a fragmented state or a rolled fragment. these observations have been summarised in appendix 5. the newcomers of molluscs from the skagen well will be mentioned. these also represent the species earlier known to live close to the recent danish waters and species new compared to what is known to be part of the recent danish fauna. this part of the record has the highest diversity and number of specimens compared to other sections of the skagen well. the mean species diversity per sample shows a rise compared to the older part and reflects the new sedimentary facies. however, the near to shore situation also puts forward the question of whether part of the faunas, if not all, may have been reworked. eliminating the uppermost ten samples covering the 5 m which can be regarded as the medieval shore. first the species represented by only few finds that are commonly found in great quantities will be discussed. lacuna pallidula occurs only as a single find at 30.0– 30.5 m level. this species occurs on fucus serratus and in great quantities from the littoral and to a depth of 70 m. hydrobia ulvae occurs normally in high numbers in shallow water. in the skagen well it has been recorded from only two levels (11.70–11.80 m and 25.0– 25.5 m) and with few specimens. rissoa violacea is connected with seaweeds and found from the tidal zone to a depth of 50 m. here the only finds are from the 27.0–27.5 and 28.0–28.5 m levels. also bittium reticulatum appears not to be part of the environment, since this species has only one occurrence at the 22.0– 22.5 m level. this species lives on zostera, as do other of the above-mentioned species. it can be concluded that the upper 30 m section of the well lacks the normal abundance of epifaunal elements connected with vegetation. this is also in good accordance with the high rate of sedimentation. geus bulletin no 3.pmd 28-06-2004, 08:45107 108 0 10 20 30 40 50 60 70 80 90 100 weight per cent, % d ep th b el ow s ur fa ce , m la bo ra to ry n um be r clay/silt medium sand gravelfine sand coarse sand 2.2 20.8 3.6 4.6 5.6 6.6 7.6 8.6 9.6 10.6 11.6 12.6 13.6 14.6 15.6 16.6 17.6 18.6 19.6 21.6 22.6 23.6 24.6 25.6 26.6 27.8 28.4 29.8 30.6 295.93 296.93 297.93 298.93 299.93 300.93 301.93 302.93 303.93 304.93 305.93 306.93 307.93 308.93 309.93 310.93 311.93 312.93 313.93 314.93 315.93 316.93 317.93 318.93 319.93 320.93 321.93 323.93 324.93 histogram of 29 grain-size analyses from the upper 30 m of the skagen well 4 among the gastropods occurring in the upper part of the well, aporrhais pespelicani occurs in the interval from 22.0–22.5 m to the 11.0–11.5 m level. this species is regarded as sublittoral from depths of 10– 180 m on a sandy muddy bottom or muddy bottom. however, it has been found in large quantities as empty shells on the shore of the east coast of skagen. this was rather puzzling until it was explained that the hermit crab might have been the actual agent bringing the shells on shore (g.h. petersen, personal communication 1998). the occurrence of lunatia montagui is restricted to the 20.0–20.5 m level, while lunatia alderi is rather frequent in the core samples. the impact of these predatory gastropods on the fauna – 15 species have been recorded with such borings, including some of the lunatia species themselves – has been quantified in appendix 3. the high number of lunatia alderi in the upper 30 m is in accordance with the preferred environment of clean sand of this species. a new neogastropod to the fauna of the well is the buccinum undatum from the 11.0–11.5 m level, while hinia pygmaea now becomes common, occurring in most of the samples from the 28.0–28.5 m level to 6.0– 6.5 m b.s. and represented in many specimens – some of them bored by predatory gastropods, as seen in appendix 3. the small gastropod oenopota turricula has a wide depth range (20–200 m), so the single finds at the 23.5– fig. 97. histogram of 29 grain-size analyses from the upper 30 m of the skagen well 4, showing two coarsening-upwards sequences. geus bulletin no 3.pmd 28-06-2004, 08:45108 109 21.0 m level most probably reflect that only in this part of the well does the clean sandy bottom occur which is preferred by oenopota turricula. of the heterogastropod newcomers in the upper section, graphis albida from the 25.0–25.5 m level can be mentioned. this species is not recorded among the recent danish molluscs (jensen & knudsen 1995). it is found sublittorally out to a 30 m depth. hemiaclis ventrosaoccursat the30.0–30.5mand11.0– 11.5 m levels, but it is recorded in recent waters at a much deeper level: 100–200 m. neither this nor the species mentioned above is recorded from danish waters. vitreolina philippii, occurring within the interval from 29.5 to 7.0 m with seven specimens, is known from the recent danish fauna and is noted as sublittoral to a depth of 200 m. this gastropod is a parasite on echinoderms, as the other eulimidae. echinoderms are still present in the material as seen from appendix 3. from the 15.0–15.5 m level, finds of chrysallida decussata occur, which is also recorded by jensen & knudsen (1995). this species occurs at the depth interval of 14–40 m. turboniella acuta has been recorded from danish waters by jensen & knudsen (1995) although rare. the occurrence of this species in the skagen well is at the 21.0–21.5 m and 20.0–20.5 m levels with, two wellpreserved specimens. among the ophistobranchs there are some fragmentary finds which have not been identified to species ratio 1:1 rightleft > 4.0> 0.5 > 2.0 > 6.5 > 8.0 size, mm 125 48 109 147 311 321 442 569 278 1129 4332 1201 113 85 166 826 98 68 59 14 0 79 75 92 39 5.0–5.5 6.0–6.5 7.0–7.5 8.0–8.5 9.0–9.5 10.0–10.5 11.0–11.5 12.0–12.5 13.0–13.5 14.0–14.5 15.0–15.5 16.0–16.5 17.0–17.5 18.0–18.5 19.0–19.5 20.0–20.5 21.0–21.5 22.0–22.5 23.0–23.5 24.0–24.5 25.0–25.5 26.0–26.5 27.0–27.5 28.0–28.5 29.0–29.5 depth below surface, m number of valves relative size distribution left/right valves ratio valves: spisula subtruncata fig. 98. size histograms for spisula subtruncata in the upper part of the subatlantic sequence, the 29.0–29.5 m level to the 5.0–5.5 m level, with ratio on left and right valves from the 29.0–29.5 m level to the 11.0–11.5 m level. geus bulletin no 3.pmd 28-06-2004, 08:45109 110 level,but species suchasretusa truncatulaandcylichna alba are found also in the upper part of the well. a fragmentary scaphopod from the 15.0–15.5 m level has not been referable to species level. among the bivalves, many are new to the already mentioned fauna from the well, and the number of specimens is for many of the species very high in comparison to what has been recorded from the older strata. of palaeotaxodonta, nucula nitidosa is found and represented all through the interval from 29.5 to 13.0 m b.s., occurring on sand bottom, which is the preferred substrate. also nucula nucleus is found within the interval from 30.5 to 8.0 m b.s. with many (13) specimens, part of them bored as the presiding species by the predatory gastropods. in the subclass pteriomorphia, species from mytiloida and pterioida such as musculus discors at the 28.0– 28.5 level and mytilus edulis in large quantities (113 specimens) are found, albeit most of the latter as juveniles. from the 28.0–29.5 m level individuals are found (with both valves). this latter species is typical in the littoral zone, but may occur at depths out to 40 m. pectinidae have been found, but all in fragments, in the interval 25.5–12.0 m b.s. ostrea edulis occurs in the interval 28.5–7.0 m b.s. – mostly as juveniles. the subclass heterodonta, from where most of the found bivalves come also includes the species most often found and characteristic of the youngest part of the marine sequence. mysella bidentata is recorded from the entire holocene, although only a few specimens are present in the early holocene. in the latest holocene as the present 30 m, 125 specimens have been found. the closely related tellimya ferruginosa occurs apart from a single find at the 61.09–61.14 m level, from the 29.5 m level where it is common up to 8.0 m b.s. both of these species have specimens bored by the predatory gastropods. tellimya ferruginosa is a commensal on echinocardium cordatum, but can also be found on its own in the sediment. mactra stultorum has been found only in the upper part of the cored section and can be seen as connected with the clean sand that is the type of bottom preferred by this species. on a suitable bottom it may be found out to a depth of 60 m. spisula subtruncata, which has a wide distribution from the littoral zone and out to a depth of 200 m, can be found both on muddy and on sandy bottoms. it dominates the uppermost part of the sequence, with 11085 specimens! in recent waters on sandy bottoms this species is one of the most common bivalves in the kattegat at depths between 20 and 30 m (jensen & spärck 1934). on the cored material from the skagen well size histograms and counts on left and right valves have been made in order to ascertain from such measurements whether the shell material is autochthonous/ parautochthonous. as seen from the figures in fig. 98, it appears that there is an even representation of the left and right valves, and the size histograms reflect a life assemblage which also might appear from the wellpreserved gracile valves. the borings counted on valves of this species make it clear that spisula subtruncata must have been the preferred victim of the predatory gastropods in this molluscan assemblage. at the 15.0–15.5 m level around 10% of the specimens are bored (2723 individuals out of which 268 have been bored). individuals (with both valves) have been found up to the 21.0–21.5 m level, where also other bivalves have been found with closed valves. however, the most even occurrence of left– right valves also at the 15.0–15.5 m level (2147–2105 respectively) may speak in favour of an autochthonous state also at this depth. the size histogram from the same level points to the same conclusion (see fig. 98). by way of the same kind of measurements it is possible to extend the possible life-assemblages up to a level of 10.0–10.5 m b.s., where the material still is present in such quantity that the measurements can be taken as bearing. investigations performed on nearshore deposits off the west coast of jylland in the agger tange area given in appendix 4 support the view that lifeassemblages can be found near to shore at depths of up to 6–7 m. almost all the ams datings in the upper part of the well have been based on spisula subtruncata, and these datings all fall within the right relative age according to their respective levels. this is not the case with the date on donax vittatus, which has also been dated within the upper 30 m interval. as shown on the dating diagram (heier-nielsen et al. 1995; appendix 3), the donax vittatus age clearly appears as an older element in a younger part of the section. however, donax vittatus will be commented upon later in the text. solenidae species often occur in the upper part of the sequence, butoften in a fragmentary state.however, phaxas pellucidus is common in the interval between the 30.5 and 20.0 m level, where it is found in several specimens in some of the samples. it lives on different bottom types from the sublittoral at a depth of 4 m out to a depth of 150 m. however, in the skagen well there is only a single occurrence at 73.39–73.44 m b.s. geus bulletin no 3.pmd 28-06-2004, 08:45110 111 one of the dominating bivalves is fabulina fabula, which only occurs within the interval 28.5–4.0 m b.s. some of the specimens have been the victims of the predatory gastropods. this species prefers a bottom type of fine sand, which might explain the interval of occurrence in the skagen well, where there are sandy layers only in the uppermost 30 m. on a suitable bottom this species goes out to a depth of 50 m. also tellina pygmaea and angulus tenuis occur in the upper part of the sequence and only there, with the exception of a single find of angulus tenuis at the 55.30–55.35 m level. this is outside the general occurrence of this shallow-water species normally found from the intertidal zone out to a depth of 10 m. donax vittatus, which is found regularly in the interval from the 27.5–4.0 m level, but often in a fragmentary state, is a typical high-energy coastal form on a sandy bottom. as already mentioned in connection with the dates, donax vittatus also occurs as an allochthonous element, which can be seen from the many rolled fragments of this sturdy shell. its occurrence out to a depth of 20 m off high-energy shores characterises in the best way the situation by the building up of the skagen spit system. the species is not found in the kattegat region and is absent from the inner part of the limfjord. gari fervensis is found only in this upper part of the skagen well from the 27.5 to 23.0 m level. accordingly, in danish waters it is known from a depth of 15–40 m on mixed bottoms and sometimes on sandy bottoms. through most of the holocene, fragments of the genus abra have been found. abra nitida, which has a wide depth distribution from the sublittoral zone out to a depth of 200 m, has been found through the last part of the holocene from the 71.89–72.00 m level to the10.70–10.80m level mostly in single specimens. this species has its maindistribution today in thedeeperparts of the skagerrak and the kattegat on muddy bottoms. a single find of a rolled fragment of arctica islandica occurs at 14.70–14.80 m, which is the only find besides the fragment from the glacial series at the 173.67– 173.85 m level. however, the washed samples have given another specimen also from the subatlantic (appendix 2). chamelea striatula is the characteristic animal of the venus community on a sandy bottom in the north sea and the kattegat. at skagen it occurred already at the 76.34–76.50 m level (late subboreal). at this depth a change of weight per cent of clay takes place (from 13.7% to 7.8%), and the fine sand component becomes the dominating grain size with a weight per cent of 54.9. from the 30 m level, chamelea striatula is more common, and specimens with connected valves occur up to 21.0–21.5 m b.s., many of them bored by predatory gastropods, as shown in appendix 3. from the point of view that also other bivalve species have been found as whole individuals up to the 20 m level, it can be regarded as the well-established limit for an autochthonous occurrence of the molluscs. however, as seen from the observations off the agger tange area given in appendix 5, there will always be an element of allochthonous material in such a high-energy coastal environment which should be taken into account also for the skagen area regarding the uppermost part of the sequence from the skagen well. a single find of timoclea ovata is also found in the upper part of the section at the 27.0–27.5 m level. this species is today found at a greater depth than chamelea striatula, but is not very numerous. within the order myoida, corbula gibba is also well represented in the upper section, with individuals found up to a level of 27.0–27.5 m b.s. this species also shows many specimens killed by predatory gastropods. corbula gibba is found in the sublittoral zone out to a depth of 250 m. at skagen its first occurrence is at 74.89–75.00 m, in the early subboreal, but it becomes common in the subatlantic and occurs in high numbers only in the last part of the subatlantic from the 30 m level, often bored. finally, two single finds of saxicavella jeffreysi and pholas dactylus occurred at the 20.0–20.5 m level. saxicavella jeffreysi is in recent danish waters not very abundant at depths between 25 and 50 m, while pholas dactylus would only be expected to be found at depths less than 10 m. pholas dactylus is a boring form found in hard substrates, which is far from the actual sediment occurring at this level in the skagen well. the fragmentary pholas dactylus can be regarded as one of the allochthonous elements that can be seen in connection with the accessory finds mentioned in appendix 3 and commented upon below. among the accessory finds the barnacles and sea urchins dominate. also fish remains are found, often in the form of vertebrae, but an otolith appears as well. other fossil remains are serpulids, bryozoans, crustaceans, and plant and insect remains, which taken as a whole very well characterise the near-shore environment. on the other hand, no concretions are found like the ones from the younger yoldia sea, or pyrite as found at the base of the holocene and the eemian. although these accessory elements cannot be quantigeus bulletin no 3.pmd 28-06-2004, 08:45111 112 fied, they offer some additional information when considered together with the sedimentological and mollusc records. conclusive remarks on the skagen well in the description given above, the faunal record is the basis for understanding the climatic changes in the skagen well, supplemented by the observation on the changes in the sediments. however, the changes found during the holocene are most likely to be connected with changes in facies, and here the changing depth is the most prominent agent, ending up with the last event represented by the depositional history of the skagen spit. based on the dating of the holocene and the late weichselian, the descriptions have been given in terms of episodes. especially the holocene strata points to a development from deeperto shallow-water facies from preboreal to subatlantic. in this development there appears to be a facies change that can be compared to the bottom communities as known from the skagerrak–kattegat region when going from the deeper-water communities of the present day, the so-called maldane-ophiura sarsi community, to the venus community of the more shallow seas. the mollusc assemblages in the skagen sequence indicate a deeper-water facies during the eemian, the weichselian and the older holocene in contrast to what hitherto was known in other parts of the danish area during the late quaternary. the skagen well has a record of the changing seas during the late quaternary, from the eemian through the weichselian (although only in parts) and the holocene. for the first time within the danish area, the full record of the marine environmental transition from the late pleistocene to the holocene can be demonstrated on the basis of molluscs. however, not all the episodes known from the skagen well can be found in marine facies of the other regions, but thanks to the new records from the north sea around the jydske rev area, a near to full holocene marine record is at hand, including part of the preboreal (petersen 1998). the environmental changes through time in the seven sectors based on the molluscan records the recorded mollusc species within each area are given in appendix 6. regarding the environmental changes through time within the danish realm, the seven sectors will be considered from the eemian, starting in the south within the classical area where forchhammer named the deposits the cyprina clay. eemian species sorted after climatic affinities the bælt sea age: eemian climatic regions: asbl class bivalvia subclass heterodonta order myoida mya truncata linnaeus 1758 total for climatic regions asb. : 1 (1.7%) climatic regions: .sb. class bivalvia subclass heterodonta order veneroida arctica islandica (linnaeus 1767) order myoida zirfaea crispata (linnaeus 1758) total for climatic regions .sb. : 2 (3.4%) climatic regions: .sbl class bivalvia subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 order pterioida geus bulletin no 3.pmd 28-06-2004, 08:45112 thermal structure of the deep lopra-1/1a borehole in the faroe islands 91 thermal structure of the deep lopra-1/1a borehole in the faroe islands niels balling, niels breiner and regin waagstein information on temperature, temperature gradients, thermal conductivity and heat flow from the c. 3.5 km deep lopra-1/1a borehole in the faroe islands is presented and analysed. the upper 2450 m of the drilled sequence consists of thick tholeiitic basalt flows and the deeper parts of hyaloclastites and thin beds of basalt. temperature data originate from high precision temperature logging a long time after drilling to a depth of 2175 m (the original lopra-1 borehole) and from commercial temperature logs measured a short time after drilling to a depth of 3430 m (lopra-1/1a). the high-precision temperature log determines accurately levels of inflow of groundwater to the borehole and significant thermal disturbances to a depth of c. 1250 m. below 1300 m, no significant disturbances are seen and interval temperature gradients for large depth intervals show only small variations between 28 and 33°c/km. the mean least-squares gradient for the depth interval of 1400–3430 m is 31.4°c/ km. in clear contrast to these overall very homogeneous, large-interval, mean temperature gradients, great local variability, between gradients of 20–25°c/km and 45°c/km, was observed between about 1300 and 2175 m (maximum depth of the high-resolution temperature log). these gradient variations are interpreted to be due to thermal conductivity variations and to reflect varying secondary mineralisation and mineral alterations. a preliminary analysis of the lopra-1/1a temperature–depth function in terms of long-term palaeoclimatic signals indicates subsurface temperatures below about 1300 m to be in equilibrium with mean surface temperatures significantly below zero during the last glacial period. a subsequent temperature increase of 12–16°c occurred at around the termination of the last glaciation. the measured temperatures (some after correction) and the thermal regime below 1300 m seem to represent conductive equilibrium conditions without significant disturbances from the effect of drilling, groundwater flow or long-term palaeoclimatic surface temperature variations. thermal conductivity measured on samples of basalt taken from drill cores and surface outcrops in the area of the borehole shows values within a rather narrow range and a well-defined mean value for low porosity basalts of about 1.8 w/m°c , while a few samples of lapilli-tuff/tuff from the borehole gave values around 1.9 w/m°c . lapilli-tuff and tuff seem to have higher matrix (grain) conductivity than basalt. heat flow is estimated at 60 ± 5 mw/m2. a heat flow of this magnitude is consistent with the faroe islands being underlain by continental crust. keywords: lopra-1/1a borehole, faroe islands, temperature gradients, thermal conductivity, heat flow ____________________________________________________________________________________________ n.b. & n.b., department of earth sciences, university of aarhus, finlandsgade 8, dk-8200 aarhus n, denmark. e-mail: niels.balling@geo.au.dk r.w., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. © geus, 2006. geological survey of denmark and greenland bulletin 9, 91–107. available at: www.geus.dk/publications/bull geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1991 92 deep boreholes generally provide the most reliable and undisturbed direct information on temperature, temperature gradients, thermal conductivity and heat flow. temperatures and temperature gradients measured in shallow boreholes may be perturbed for a variety of reasons including local effects of groundwater movements, topography and shortand long-term palaeoclimatic surface temperature variations. the deep lopra-1/1a borehole provides a unique opportunity for obtaining deep thermal information from an area of the north atlantic that is of considerable interest to both the general geoscience community and to the hydrocarbon industry. the lopra-1/1a borehole is situated in the southern island (suðuroy) of the faroe islands (fig. 1) (at 61°26′ 36″n, 6°46′30″e). it was drilled in 1981 as a research borehole to a depth of 2175 m below ground level (berthelsen et al. 1984). in 1996 the borehole was re-entered by a consortium of exploration companies and deepened to 3565 m measured depth corresponding to a vertical depth below ground level of 3540 m. all depth values given in this paper (if not stated otherwise) are vertical depths measured from ground level 8.8 m above mean sea level. thermal measuring results from the original lopra-1 borehole were presented by balling et al. (1984). the purpose of the present paper is to integrate, analyse and discuss all available thermal information from the whole depth range of the lopra-1/1a borehole. it includes new high-precision continuous temperature logging results from the original hole measured a long time after drilling, temperature measurements from the deepened part acquired as part of the commercial logging runs during and shortly after drilling. new thermal conductivity measurements from core material from the deepened section and from surface exposures in the lopra-1/1a area have also been made. temperatures and temperature gradient variations are analysed in relation to disturbances from groundwater flow, variations in rock thermal conductivity and information on long-term palaeoclimatic surface temperature variations. a new terrestrial heat-flow value for the lopra-1/1a site is presented. geological environment and lithology the volcanic succession of the faroe islands the faroe islands form part of the palaeogene north atlantic province of tholeiitic flood basalts. the faroe volcanic succession has been divided informally into the upper, fig. 1. a: location map of the faroe islands in the northern north atlantic. b: the lopra-1/1a borehole was drilled in the southern faroese island of suðuroy. the total thickness of volcanic sequences in the faroe islands is at least 6.5 km. about 3 km is exposed, and 3.54 km was drilled at lopra-1/1a. dashed lines a and b mark boundaries between lower and middle and middle and upper basalt series respectively. 3 2 3 2 km m id dl e se ri es ex po se d dr ill ed up pe r se ri es lo w er s er ie s b a b a a b 1 1 0 3.5b b a faroe islands sandoy suðuroy lopra streymoy faroe islands 10 00 m scotland10 00 m 10 00 m 10 00 m 10 00 m 10 00 m 60° 20° 20° 60° 10° 0° 70° 20° 10° 0° jan mayengreenland 70° iceland geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1992 93 middle and lower basalt series or formations with a total thickness of more than 6.5 km (rasmussen & noe-nygaard 1970; waagstein 1988). the lowermost 3.5 km, which is entirely in the lower basalt formation, is known only from the lopra-1/1a borehole (fig. 1). both seismic and other geophysical evidence (bott et al. 1974; richardson et al. 1998) and geochemical data (garièpy et al. 1983; hald & waagstein 1983; holm et al. 2001) indicate that the faroe islands are underlain by continental crust. pre-volcanic rocks have not been reached by drilling. linear magnetic anomalies associated with oceanic crust occur 60–70 km north of the islands (skogseid et al. 2000). the flood basalts were formed by extensive volcanism associated with the continental splitting between nw europe and east greenland in upper paleocene to lowermost eocene time (e.g. skogseid et al. 2000). larsen et al. (1999) used geochemical analyses and stratigraphic correlations between the volcanic successions in the faroe islands and east greenland to interpret the faroese lower basalt formation as a pre-breakup sequence and the middle and upper basalt formations as syn-breakup sequences. since deposition of the upper basalt formation, little or no deposition has occurred in the faroes. volcanic activity continued, however, on the greenland side of the rift with the eruption of an additional 3–3.5 km of basalts in an area then located close to the centre of the icelandic mantle plume. lithology the upper 2450 m of the lopra-1/1a borehole consists of subaerial lava flows of tholeiitic basalt with an average thickness of about 20 m. most of the flows have a massive core and a vesicular rubbly top. the lavas are commonly separated by palaeosols made up of volcanic ash or material eroded from the flow tops. the sediments range from a few centimetres to more than 4 m in thickness. the deeper part of the well, from about 2450 m to total depth, consists of hyaloclastites (lapilli-tuff and tuff ) and thin beds of basalt. since deposition, secondary mineralisation and mineral alterations have occurred. the bulk thermal properties of the basaltic sequences seem to be controlled mainly by the two major minerals feldspar and pyroxene, which occur in roughly equal amounts in common basalts. however, it appears from our thermal gradient analysis that within-flow variations of the degree of alteration of the basalt is an important controlling factor for local variations in rock thermal properties and hence temperature gradient variability. some information on secondary mineral alterations and mineralisation is thus required for a proper thermal analysis. olivine is a minor constituent that has been generally replaced by clay. haematite has formed from iron-rich minerals under oxidising conditions, especially within flow tops and interbasaltic sediments. the original plagioclase feldspar is partly or completely replaced by albite in the deepest part of the borehole due to very low-grade burial metamorphism. originally variable amounts of interstitial glass representing frozen melt are completely altered to clay and other secondary minerals. most gas vesicles and pores and fractures once filled with free water are now, particularly at great depth, partly or completely filled with low-temperature minerals deposited by flowing groundwater. these mineralisations consist dominantly of clay and zeolite minerals, silica minerals (chalcedony, agate, quartz) or calcite. the vertical distribution of secondary minerals and zeolite zones of the lopra-1/1a borehole are described by jørgensen (1984, 2006, this volume). the lithology and chemistry of the upper 2.2 km sequence is described in detail by hald & waagstein (1984). temperature and temperature gradients temperature logs and conditions of measurements temperature information is available from several temperature logs. measurements have been carried out by several techniques, either as dedicated temperature logging or in combination with other log operations. temperatures were measured both during interruptions in the drilling and after the drilling was completed. this applies both to the original lopra-1 borehole and to the deepened part of lopra-1/1a. temperature logs were run in the original lopra-1 borehole by the icelandic energy authority, orkustofnun, the operator of the original hole. the last one was run in 1983, 17 months after drilling. these results are described in detail in balling et al. (1984). a more detailed continuous temperature log has since been run in the original lopra-1 hole. it was carried out by the present research group in 1994 to a depth of 2175 km, almost 13 years after drilling, using the high-precision quartz-oscillator system of the university of aarhus. measurements were taken while running down-hole at a nominal speed of 10 cm/s with a sample rate of two seconds resulting in a sample interval of about 20 cm. relative temperature resolugeus bulletin no 9 7 juli.pmd 07-07-2006, 14:1993 94 tion is better than 0.005°c and absolute accuracy is calibrated to about 0.05°c. all temperature logs from the lopra-1/1a deepened section below 2175 m were acquired by the company schlumberger in combination with other logging operations relatively soon after drilling activities and circulation of drilling fluid. the temperature data are available as standard six-inch point measurements taken by thermistor probes and are estimated to have an accuracy better than 0.1°c. several temperature logs are thus available both for the original lopra-1 hole and the deepened lopra-1/1a section. five of these have been selected as those giving the most valuable information for interpretation (table 1 and fig. 2). they cover depths from the surface to 3430 m. drilling and circulation of drilling fluid disturb significantly the temperature structure of the borehole and the unperturbed so-called equilibrium temperature–depth distribution can be measured only a relatively long time after drilling. if sufficient time has not passed, corrections to measurements must be applied (cf. beck & balling 1988). in general, during circulation of drilling fluid, the upper part of the hole is heated and the lower part is cooled. the time needed for a borehole to reach temperature equilibrium depends on several factors including drilling history, temperature of the drilling mud and the required accuracy of temperature and temperature gradients, but may be relatively long compared to the duration of the drilling. for deep boreholes like the lopra-1/1a borehole, at least one to two years may be needed to obtain both accurate equilibrium temperatures over the whole section and accurate local temperature gradients. only for the bottom part of the hole and the neutral zone of least disturbances may near-equilibrium temperatures be measured much sooner after drilling and last drilling fluid circulation. as mentioned above, temperature measurements were carried out in the original lopra-1 borehole to a depth of 2175 m a long time (up to almost 13 years) after drilling. the temperature logging results from the new section fig. 2. selected measured temperature–depth profiles. logs 1 and 2 are from the original borehole lopra-1 and logs 3, 4 and 5 were measured in the deepened part. information on time of measurements, details of depth intervals and time after drilling or last drilling fluid circulation are given in table 1. log temperatures presented here are uncorrected. corrected temperatures are given in fig. 3. temp. log 1 2 3 4 5 date of measurement 23.03.1983 04.08.1994 01.10.1996 02.10.1996 30.10.1996 depth of borehole (m) 2175 2175 3144 3144 3496 measuring agency orkustofnun, reykjavik university of aarhus schlumberger, esbjerg schlumberger, esbjerg schlumberger, esbjerg 0–1974 0–2175 3020–3095 2170–3075 2990–3430 time after drilling/circulation† 17 months 12 years 9 months 27 hours 50 hours 53 hours depth interval of temperature data (m) * table 1. basic information about temperature logs from the lopra-1/1a borehole * † five logs have been selected to give the most valuable temperature information. the time elapsed between the last significant disturbance from drilling or last circulation of drilling fluid and temperature logging. 400 0 800 1200 1600 2000 2400 2800 3200 3600 d ep th ( m ) 5 3 4 2 21 1 −10 0 20 40 60 80 100 temperature (°c) geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1994 95 drilled in 1996 were, however, carried out no more than 27–53 hours after last drilling fluid circulation, cf. table 1, and corrections must be considered. temperature disturbances are created not only by the process of drilling. the upper part of the borehole is also affected by upward water flow inside the borehole. after drilling, the lopra-1 hole (total depth 2175 m) started to flow at a rate of about 10 l/min when the drilling mud was replaced by fresh water of lower density. the well head was closed between the end of drilling and the time of log number 1 (17 months after drilling, cf. table 1) preventing water from flowing freely to the surface. the hole was opened on the day of temperature logging and after one hour it began to flow at a rate of about 10 l/min. temperature measurements were carried out from 3 to 6 hours after opening. at the time of log number 2, measured in 1994, the hole was flowing freely at about the same rate and had not been closed for several years. the highest recorded temperature of 98.6°c was measured at a depth of 3430 m, 53 hours after circulation when the hole was 3496 m deep (log 5 from 30 october 1996). the deepest point of temperature information is 3527 m where 92°c was measured on 2 november 1996, 17 hours after the latest drilling fluid circulation. two days later, when the drilling had reached its final vertical depth of 3540 m (3565 m measured depth below rotary table), a temperature of 91°c was measured at 3507 m, 22 hours after circulation of drilling fluid. these lower temperatures measured later at slightly deeper levels demonstrate the effect of cooling by drilling fluid circulation. correction of temperatures a comparison of the raw temperature data of logs 1 (measured in 1983) and 2 (measured in 1994) shows a difference of 2–3°c at depths below 400 m and an almost constant offset of 2°c between 1000 and 1600 m. such an almost constant difference is very unlikely to be caused by water flow or any other effect associated with the borehole and, from further data analysis, this difference is ascribed to an instrumental calibration offset in log 1 by about 2.0°c. (equipment used for log 2 measurements was carefully calibrated before and after logging.) after adding 2.0°c to the original log 1 values, log 1 and log 2 measurements agree to within ± 0.2°c between 1000 and 1600 m, increasing to a maximum difference of 0.7°c at 1974 m, the maximum depth of log 1. in the topmost part of the borehole, log 2 shows slightly higher temperature differences (by up to 2–3°c) due to a longer time of temperature disturbance from up-hole water flow (cf. fig. 2). measured log 1 and log 2 temperatures are both clearly elevated in the upper part of the hole because of water flow. at near-surface level, measured temperatures (fig. 2) are well above the mean ground temperature of about 7°c. by temperature gradient analysis (se below), levels of water inflow have been localised accurately. below a depth of about 1250 m, measured temperatures of logs 1 and 2 are unlikely to be disturbed significantly by flow of water in the borehole, and log 2 temperatures are assumed to represent undisturbed equilibrium values. the temperatures on logs 3, 4 and 5, from depths between 2170 and 3430 m (table 1 and fig. 2), were measured between 27 hours (log 3) and 53 hours (log 5) after drilling fluid circulation following drilling activities and are thus disturbed. measured temperatures on logs 3, 4 and 5 are, due to their deep position in the hole, lowered 400 0 800 1200 1600 2000 2400 2800 3200 3600 d ep th ( m ) 5 4 2 2 −10 0 20 40 60 80 100 temperature (°c) fig. 3. measured temperatures on log 2 and corrected temperatures of logs 4 and 5. corrections were applied for the estimated effect of drilling and drilling fluid circulation. the dashed line has a constant gradient of 31.4°c/km calculated by least-squares for the depth interval of 1400–3450 m and extrapolated to the surface. the upper part of the borehole (above c. 1200 m) is disturbed by upward flow of water inside the borehole. the negative temperature at the surface intercept indicates that temperatures below about 1200–1300 m are in equilibrium with a palaeosurface-temperature significantly below that of the present-day mean surface temperature of about 7°c. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1995 96 by the circulation of drilling fluid to temperatures below formation temperature. they are thus all lower than the undisturbed formation equilibrium values. the depth intervals over which measurements were made at different times overlap partly. this makes it possible to estimate the size of temperature disturbances and correct for them. the upper part of log 4 overlaps with the lowest part of log 2 in the depth interval 2170–2175 m. log 4 temperatures were here 3.7°c below the temperatures of log 2, which are assumed to be undisturbed. the increase in temperature between logs 3, 4 and 5 (fig. 2), combined with additional log data not shown, has been used to estimate the amount of disturbance by hornertype analysis. in the deepest part of log 5 a temperature of 98.6°c was measured at a depth of 3430 m, 53 hours after drilling fluid circulation. temperatures are estimated to have been reduced by 3–7%, so, applying a correction of 5%, the undisturbed value is about 103.5°c. in the final selection of temperature data, only the almost linear part of log 4 (2170–2775 m, fig. 2) with corrections between 3.7°c (top) and 4.0°c (bottom) was used. after corrections of logs 4 and 5 for the estimated disturbance due to drilling, the corrected temperatures follow the same depth trend as that of the deeper part of log 2. below 1100 m, measured temperatures on log 2 and the corrected values on logs 4 and 5 fall within 1.2°c of a constant gradient least-squares temperature line (fig. 3). as discussed above, the corrected temperatures on logs 4 and 5 are thought to represent equilibrium temperatures to a good approximation (within 1–3°c) and to be sufficiently accurate to calculate accurate mean temperature gradients for the larger depth intervals. measured temperatures on log 2 (below 1200 m) and corrected values on logs 4 and 5 are listed at 100 m depth intervals in table 2. equilibrium temperatures and temperature gradients mean least-squares temperature gradients from selected logs and depth intervals have been calculated and are listed in table 3. in order to ensure a homogeneous base for the calculation of temperature gradients, all logs were resampled to depth increments of 5 m. mean gradients show only small variations between 28 and 33°c/km. log 2 yields a mean gradient of 32.9°c/km between 1400 and 2175 m and the combined data from logs 2, 4 and 5 for the depth interval 1400 to 3430 m yield a temperature gradient of 31.4°c/km. this demonstrates an overall very homogeneous thermal gradient structure. in clear contrast to the above overall small gradient variations, significant local temperature gradient variability is observed. figure 4 shows running mean least-squares interval temperature gradients (covering 5, 25 and 100 m depth intervals) derived from the high-resolution log 2 run from surface to 2175 m. table 2. listing of selected temperatures from logs 2, 4 and 5 1200 1300 1400 1500 1600 1700 1800 1900 2000 2100 2200 2300 2400 2500 2600 2700 2800 2900 3000 3100 3200 3300 3400 3430 66.2 69.1 72.0 74.7 77.8 81.5 n.d. n.d. 90.5 93.7 96.4 98.9 102.4 103.5 34.1 36.2 39.4 42.6 46.1 49.4 52.3 55.6 58.8 62.5 temperatures (oc)* depth (m) log 2† log 4‡ log 5‡ measured temperatures above 1200 m are disturbed by water flow inside the borehole (cf. fig. 3) and are not listed. temperatures of log 2 are assumed to represent accurately the undisturbed equilibrium temperatures. temperatures of logs 4 and 5 are corrected for the estimated influence of drilling disturbances and may represent equilibrium temperatures to within ± 1–3°c. n.d.: no data. ‡ † * table 3. least-squares mean temperature gradients for various depth intervals depth interval (m) 1400–2175 2170–2770 2995–3430 1400–3430 temp. gradient (oc/km) –6.7 0.0 6.2 –3.9 log 2 4 5 2, 4, 5 32.9 30.1 28.2 31.4 * intercept temperature value (linearly extrapolated temperature at *intercept (oc) zero depth) associated with each depth section. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1996 97 fig. 4. running mean least-squares temperature gradients for 5, 25 and 100 m depth intervals as indicated. temperature gradients were taken from the high resolution log 2. the dashed line shows the assumed unperturbed mean background gradient of 32.9°c/km calculated for the depth interval of 1400–2175 m (cf. table 3). levels of significant inflow of groundwater into the upper part of the borehole are clearly seen as local gradient maxima. below about 1250 m, temperatures and temperature gradients are thought to represent generally conductive equilibrium conditions and gradient variability is mostly due to variations in rock thermal conductivity (see also fig. 5). d ep th ( m ) 0 20 40 60 200 400 600 800 1000 1200d ep th ( m ) 1400 1600 1800 2000 2200 5 m 0 20 40 60 25 m 0 20 40 100 m temperature gradient (°c/km) 0 the upper part of the log is disturbed by water flow. the original hole was uncased below 180 m and water at above hydrostatic pressure was able to enter the hole through local fractures or permeable beds. the temperature and temperature gradient logs combined show clearly levels of significant disturbance due to inflow of water to the borehole. they are characterised by a locally high temperature gradient. above the level of inflow, both temperature gradients and temperatures are reduced. at the approximate depths of 292, 360, 444, 1111, 1132 and 2130 m, temperature drops of between 0.2 and 0.6°c are observed, resulting in locally high temperature gradients. the highest local temperature anomalies of 0.5–0.6°c occur at 444 and 1132 m, where also maximal temperature gradient perturbations are observed (fig. 4). a local minor disturbance is seen around 1538 m (see also fig. 5). most of these thermal perturbations, in particular those above 1100–1200 m, are easily interpreted in terms of inflow of ground water at approximately local formation temperature into a section of the borehole where temperatures are artificially elevated due to upwards-flowing water coming from deeper levels of higher formation temperature. local lowering of borehole temperature may also occur if water flows downwards through inclined fractures from levels of lower formation temperature to levels of higher formation temperature. however, this does not seem to have happened here. longer wavelength temperature-gradient minima are observed at about 550–800 m and 1160–1250 m that are not clearly associated with localised zones of inflow of water (fig. 4). these zones of low gradient may be due to broader zones of water flow into the borehole, perhaps associated with an almost steady migration of ground water within porous or fractured parts of the formation that was initiated long before drilling. below 1200–1300 m, the temperature–depth function of log 2 follows the overall almost linear trend of temperature increase with depth (fig. 3). this trend is overprinted by significant local temperature gradient variations down to the maximum depth of log 2 of 2175m (fig. 4). the local gradient variations are of a different character from those discussed above, but might at first geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1997 98 volcanic flows division 0 20 40 60 temperature gradient (°c/km) 101418 calliper (inch) 2.0 2.4 2.8 density (103 kg/m3) 020 1.6 1.2 0.840 apparent porosity (%) colour red2/ (green x blue) 1200 1300 1400 1500 1600 1700 1800 1900 2000 2100 d ep th ( m ) fig. 5. running 5 m mean temperature gradients from log 2 shown together with the calliper, density and neutron porosity logs over the depth interval 1200–2170 m. the left-hand column shows the interpreted volcanic flow boundaries (hald & waagstein 1984) with closely spaced lines generally indicating thin beds of sediment and the right-hand column shows the relative intensity of red measured from cuttings. see text for further details. note the reversed scale for calliper, apparent porosity and colour. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1998 99 sight indicate similar disturbances from flow of water. however, a detailed comparison between temperature gradient variability and other petrophysical log characteristics, including density and neutron porosity logs as well as calliper log data, shows remarkable correlations beginning at a depth of 1250–1300 m (fig. 5). mean least-squares temperature gradients based on 5 m averaging intervals reveal local maximum gradients of up to 40–45°c/km and local minimum values down to 20–25°c/km. local intervals of high temperature gradient correlate with intervals of high density, low neutron porosity and decreased calliper. intervals of low temperature gradient correlate with intervals of low density, high neutron porosity and increased calliper. the correlation between temperature gradient variability and short-range variations in borehole calliper is particularly remarkable. maximum temperature gradient variability and close correlations are most pronounced around 1450–1550 m and 1750–1850 m, but a good general correlation with the physical properties of the various basaltic flow units is observed for most of the depth range shown in fig. 5. a correlation of low calliper with increased temperature gradients cannot be explained by potential water flow inside the borehole at these greater depths. any flow of water would have the opposite effect of lowering gradients in narrow parts of the borehole due to locally increased flow rates. above a depth of about 1200 m no correlation is observed. here temperature gradients seem to be controlled by the flow of water. the temperature gradient variations below about 1250 m consequently need to be explained in terms of a linkage between resistance to the drill bit, the mineralogical composition and structure and bulk rock thermal properties of the formation. these observed correlations and the inferred variations in rock thermal conductivity are discussed in detail below. influence of palaeoclimate surface temperature variations penetrate into the subsurface, and temperature and temperature gradient measurements from boreholes may be used to extract information on short-term as well as long-term surface palaeoclimatic temperature variations (e.g. dahl-jensen et al. 1998; huang et al. 2000; kukkonen & joeleht 2003). the lopra-1/1a borehole is not particularly well-suited for such purpose due to disturbances of the upper part of the hole by water flow. we shall not, therefore, go into detail, but deal only with some main effects of long-term climatic variations. simple linear extrapolation of the temperature–depth function from the deeper parts of the borehole to the surface yields a negative intercept temperature of –3.9°c for the depth interval of 1400–3430 m and of –6.7°c for the interval 1400–2175 m (fig. 3 and table 3). the low intercept temperatures indicate that temperatures in the deeper parts of the borehole are in equilibrium with a surface temperature significantly below the present-day mean ground temperature of about 7°c. this simple linear extrapolation of deep mean temperature gradients seems justified in our preliminary study because of the apparent homogeneity of mean thermal conductivity and temperature gradients over large depth intervals (see also next section). the shorter intervals of 2175–2770 m and 2995– 3430 m give higher intercept values, but these are considered uncertain because of the short length of the intervals that make the extrapolation more sensitive to local variations in thermal conductivity. forward thermal modelling using a thermal diffusivity of basalt of 0.7 × 10–6 m2/s shows that long-term surface temperature variations of the magnitude associated with glaciation and deglaciation are reduced in amplitude to less than 1–2°c (depending on surface temperature amplitudes) at depths below 1300–1500 m. the most accurate temperature gradient and temperature intercept value is probably that from the depth interval 1400–2175 m. since porosity is likely to decrease with depth, porosities in the rocks sampled by the upper part of the borehole may on average be slightly greater than in those sampled by the deeper parts. this could point towards a slightly lower thermal conductivity and hence a slightly higher temperature gradient at shallow depths. this effect may, however, be more or less cancelled by secondary mineralisation, which tends to increase the conductivity of porous sections. this means that the mean characteristic conductivity of the upper 1400 m may not differ much from that of the interval 1400–2175 m, justifying the extrapolation with a constant gradient. we may thus interpret the surface intercept temperature of –6.7°c as an estimate of the long-term characteristic mean for the cold period of the last glaciation. we estimate the increase in surface temperature associated with the termination of the last glaciation to be of the order of 12–16°c. this is a preliminary estimate considering the approximate nature of our procedure, uncertainties related to the lack of good temperature data from the upper part of the borehole (which prevents extraction of a detailed past temperature-time function) and uncertainties related to possible vertical variations in thermal conductivity. a more detailed analysis of this problem, including inverse modelling, must be based on thermal information from other boreholes, in particular from near-surface ingeus bulletin no 9 7 juli.pmd 07-07-2006, 14:1999 100 tervals where the lopra-1/1a borehole is disturbed thermally. this is beyond the scope of this paper. from inverse analyses of the temperature–depth function from the grip-borehole on the greenland ice sheet, dahl-jensen et al. (1998) calculated a surface-temperature increase of 23°c following the last glacial maximum. using a similar procedure on many boreholes, kukkonen & joeleht (2003) obtained an average warming of 8°c for nw europe at the termination of the last glaciation. the deviation in the upper part of the lopra-1/1a borehole from the general linear trend of temperatures deeper down (fig. 3) thus has two main causes. the increase at surface level from about 7°c to the measured borehole temperature at surface of about 12°c (log 2) is due to the upward flow of warm water inside the borehole. the remaining part, the increase from negative intercept values to about 7°c, is interpreted to reflect the increase of temperature at around the termination of the last glacial period. the penetration of this heating effect to a depth of about 1200 m (fig. 3) is in agreement with model calculations. unfortunately, this depth level is also the approximate level above which temperatures and temperature gradients are significantly disturbed from the flow of water inside the borehole. thermal conductivity a limited amount of suitable sample material was available for thermal conductivity measurements. only one conventional core was drilled within the deepened section below 2175 m. however, a few rotary sidewall cores were also long enough to be measured. including published measurements on four cored sections from the original borehole, a total of 11 samples have been measured covering the depth range of 337 to 3531 m (driller’s depths). rock materials measured represent massive basalt (7 samples), lapilli-tuff (3 samples) and tuff (1 sample). in addition to thermal conductivity, rock density and porosity were also measured. measuring results on samples from the lopra-1/1a borehole are summarised in table 4. supplementary preliminary conductivity measurements were additionally carried out on 14 samples taken from surface exposures of basalts near the borehole. measuring techniques both the needle-probe transient line source technique and the steady-state divided bar technique were applied to measure thermal conductivity. these are standard techniques table 4. measured thermal conductivity, porosity and density and calculated matrix (grain) thermal conductivity and density c1† c2 c3 c4 c5 swc 46‡ swc 37 swc 13 swc 6 swc 5 swc 4 mean values, all samples (11) standard deviation mean values, basalt (7) standard deviation mean values, lapilli-tuff/tuff (except outlier swc 37) (3) standard deviation sample basalt basalt basalt basalt basalt basalt lapilli-tuff lapilli-tuff tuff lapilli-tuff basalt 337.5 860.1 1218.1 2177.3 2380.0 2441.0 2562.0 3438.0 3512.5 3514.5 3531.0 depth* (m) 1.75 1.85 1.74 1.79 1.79 1.87 1.35 1.88 1.91 1.93 1.84 1.79 0.16 1.80 0.05 1.91 0.03 thermal conductivity (w/moc) 1.81 1.95 1.80 1.82 1.85 1.96 1.60 2.29 2.17 2.27 1.90 1.95 0.21 1.87 0.07 2.24 0.06 matrix thermal conductivity (w/moc) 2.96 2.98 2.94 3.00 3.00 2.93 2.47 2.76 2.78 2.68 3.10 2.87 0.18 2.99 0.06 2.74 0.06 bulk density (103kg/m3) 3.02 3.07 3.00 3.03 3.06 3.01 2.78 3.07 2.98 2.91 3.16 3.01 0.10 3.05 0.05 2.99 0.08 matrix density (103kg/m3) 3.0 4.5 3.0 1.7 2.7 3.9 17.3 14.8 9.9 12.1 2.7 6.9 5.6 3.1 0.9 12.3 2.5 porosity (%) rock type * driller’s depth below rotary table (to top of core). † samples c1 to c5 are from conventional cores taken during drilling. ‡ the swc samples are small rotary sidewall cores taken after drilling. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19100 101 (l–ϕ) for laboratory rock thermal conductivity measurements (e.g. beck 1988). all measurements were carried out in the geophysical laboratories of the university of aarhus. equipment and measuring procedures were similar to those described in balling et al. (1981). most rock materials from the deepened section of the borehole and from surface exposures were measured by the needle probe technique. the needle probes used have a nominal length of 50 mm and an outer diameter of 1.5 mm. the measured sidewall cores are cylindrical with a diameter of about 24 mm and lengths ranging from 18 to 30 mm. the probe length cannot be reduced significantly since interpretation is based on line source approximations. the probe length thus exceeds that of the samples to be measured. this difficulty was largely overcome by placing the rock sample of unknown conductivity along the critical central position of the needle probe where the temperature rise function is measured and extending the sample by ‘end materials’ of known conductivity close to that of the material to be measured. by iterative trial and error procedure, the difference between conductivity of end materials and conductivity of rock sample was reduced to less than 0.3 w/m °c. experience suggests that this is sufficiently small for an accurate sample conductivity measurement. any further boundary effects due to the small size of samples were minimised by immersing the samples in water-saturated sand with conductivity close to that of the rock samples being measured. all samples were water-saturated under vacuum before measurement, which was carried out at normal laboratory temperature (about 20°c) and pressure (1 atm.) conditions. all needle probes were calibrated and tested by measurement of standard materials of known thermal conductivity. the heating period for the probes was 40–60 seconds and the temperature rise at probe centre typically 2–4°c. sample conductivity was determined as a mean value of at least three individual measurements and individual measurements on the same sample generally did not differ by more than 3–5%. the unknown thermal conductivity was calculated from the temperature rise data using the iterative least-squares inversion technique of kristiansen (1991). samples c1 to c5 (table 4) were originally measured by the divided bar technique and the results were reported in balling et al. (1984). these older measurements seem to be somewhat too low. a comparison of new needle probe measurements with old divided bar measuring results on material from the same basalt cores shows that the old measurements are systematically about 15% too low. the older low results seem due to the small dimensions of the samples. previous measurements are thus corrected by + 15%. considering all sources of experimental uncertainty, reported conductivity values are estimated to be accurate to ± 0.1 w/m °c. rock bulk density of water saturated samples and porosity were measured on all samples. density was measured by using the archimedes principle of buoyancy. weight of samples in air and immersed in water, respectively, yields sample weight and volume. porosity was determined by measuring loss of weight of the water-saturated samples when drying them at about 110°c for 1–2 days. repeated determinations of porosity on selected samples suggest a precision of ± 1% for low porosity samples (2–5%) and ± 1–2% for samples of higher porosity (10–20%). with known porosity, bulk density and bulk thermal conductivity, the solid matrix (grain) density and thermal conductivity may be estimated. the computation of matrix density is straightforward assuming proportional contribution of solid matrix and water. matrix thermal conductivity was computed from the geometric mean formula relating bulk conductivity, kb, matrix conductivity, km, conductivity of water, kw, and porosity (pore fraction), ϕ, by kb = k m k w . bulk conductivity and porosity are measured. the conductivity of water at room temperature is 0.6 w/m °c. measuring results thermal conductivity measuring results are listed together with the density and porosity determinations in table 4. mean values and standard deviations were calculated for all 11 samples, for the basalts alone (seven samples) and for lapilli-tuff/tuff alone (three samples, excluding the sample swc 37). variations between samples are small. all basalts are of low porosity (< 5%) and the conductivity is within the range of 1.75–1.87 w/m °c with a well defined mean value of 1.80 w/m °c. the lapilli-tuff and tuff samples have higher porosity (10–17%) and show a wider range of conductivity, between 1.35 and 1.93 w/m °c. except for sample swc 37, the lapilli-tuff/tuff samples have a higher solid matrix conductivity than the basalts. omitting swc 37, mean matrix conductivity of lapilli-tuff and tuff is 2.24 w/m °c as compared to 1.87 w/m °c for basalt. this difference in matrix conductivity explains why lapillituff and tuff have slightly higher conductivity than basalt despite a higher porosity and higher content of free water of low conductivity. the anomalous sample swc 37 has the highest porosity (17.3%) and lowest bulk and matrix conductivity measured. this is possibly due to the presence of a significant amount of secondary analcite, a mineral of very low conductivity (1.3 w/m °c) (horai 1971). ϕ geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19101 102 the thermal conductivity measurements on core materials from the lopra-1/1a borehole have been supplemented by preliminary measurements of samples of basalts taken from surface exposures in the local area of the borehole. a total of 14 samples were measured. measurements were again carried out on water-saturated samples using the needle probe technique. eight samples of low porosity (≤ 4%) have a mean conductivity of 1.77 w/m °c (range 1.67–1.86 w/m °c) and six samples of higher porosity (porosity range 5–26% and mean porosity 13%) have a mean of 1.51 w/m °c (range 1.38–1.59 w/m °c). the conductivity of low-porosity basalts is thus well defined, having a value of about 1.8 ± 0.1 w/m °c. conductivity decreases with increasing porosity due to the presence of water of low conductivity. details of these measurements are not shown, but the results conform well to and supplement those from the borehole samples. other studies our conductivity results on basalts agree well with results obtained by others. oxburgh & agrell (1982) measured more than one hundred samples of basaltic flows, intrusions and breccias covering the full depth range of the 2 km deep reydarfjordur borehole in eastern iceland. their measurements on water-saturated samples show increasing thermal conductivity with decreasing porosity and increasing sample depth. single measurements from basaltic flows and intrusions range between 1.4 and 2.2 w/m °c. mean values over 500 m intervals increased with depth from about 1.6 in the upper part of the borehole to 1.7– 1.9 w/m °c in the central and deeper part of the hole. measurements on 17 rock samples classified as breccias showed the highest values of conductivity and the widest spread, 1.6–2.8 w/m °c. as part of heat-flow measurements in shallow boreholes in the south-eastern part of the deccan volcanic province, central india, roy & rao (1999) measured thermal conductivity on about 25 core samples of basalt and several samples of fresh massive basalt from outcrops. they obtained sample values within the narrow range of 1.6– 1.8 w/m °c with a well-defined mean value of about 1.7 w/m °c. the same range of measured thermal conductivity is found in the large dataset of robertson & peck (1974) on basalts from hawaii for water-saturated samples of low porosity (2% to about 10%) and low olivine content (0–5%). with increasing porosity and pore-water content, conductivity decreased significantly and variations in mineral content played an important role (see also horai 1991). conductivity variations at lopra-1/1a and their causes in a conductive steady-state geothermal regime, variations of the temperature gradient are related to variations in rock thermal conductivity. intervals of high thermal conductivity result in low temperature gradients and intervals of low thermal conductivity result in high temperature gradients. this simple inverse relation follows from fourier’s law of heat conduction, which relates heat flow to the product of thermal conductivity and temperature gradient. for depth sections of low heat production, conductive heat flow is almost constant and temperature gradient variations will reflect variations in rock thermal conductivity. water has a thermal conductivity of 0.6 w/m °c, which is much less than a rock matrix of overall basaltic composition with a mean conductivity around 2 w/m °c. significant variations in porosity of the basalt will therefore result in major conductivity variations and associated variations in conductive temperature gradient. thermal conductivity will decrease with increased content of pore water and the temperature gradient will increase. we observe that sections of high neutron porosity are intervals of local low temperature gradient (fig. 5). this means that a significant part of the water in rocks of apparent high porosity must be bound in water-bearing secondary minerals. furthermore, some of the secondary minerals (water-free or not) must have a thermal conductivity significantly above that of the mean value of the minerals of unaltered basalt. maximum temperature gradients are generally observed within the massive (non-porous) cores of basalt flows, which are characterised by high density and low neutron porosity (fig. 5). this mainly reflects the presence of feldspar, a primary igneous minerals of low conductivity. however, secondary minerals of even lower conductivity must be present as well. the large local conductive temperature gradient variations observed between about 1250 and 2175 m (figs 4, 5) are thus interpreted to originate from significant vertical variations in mean thermal conductivity. in order to maintain a constant heat flow of around 60 mw/m2 (see next section), local intervals of minimum temperature gradient of 20–25°c/km must indicate conductivities within the approximate range of 2.5–3.0 w/m °c, and local intervals of maximum temperature gradients of about 45°c/ km must indicate conductivities around 1.3–1.4 w/m °c. the thickness of lithological units of maximum temperature gradient variations and inferred maximum conductivity variations is typically in the range of 5–20 m (fig. 5). geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19102 103 such variations of rock thermal conductivity by a factor of about two are not directly represented in our set of conductivity measurements (table 4). a potential for variation is, however, indicated by the observation that the mean solid matrix conductivity of lapilli-tuff and tuff is 20% higher than that of basalt. zones of inferred increase of conductivity are observed to be closely related to zones of reduced density in the originally porous part of the basalt flows. this may be explained by the presence of secondary minerals of high conductivity. the secondary filling of pores, voids and cracks must include minerals of thermal conductivity significantly above that of normal basalt matrix with a conductivity around 2 w/m °c. this interpretation is consistent with observations of oxburgh & agrell (1982) who found that thermal conductivity in the reydarfjordur borehole generally increased with the degree of alteration, with the highest conductivity of up to 2.8 w/m °c occurring in rock samples broadly classified as breccias. in general, the thin sediment intervals occur within the broader intervals of low temperature gradient (fig. 5), indicating a mean thermal conductivity of the sediments very close to that of the adjacent basalt flows. values of thermal conductivity quoted below are mostly from the comprehensive study and listing of conductivity of rock forming minerals by horai (1971). high-conductivity secondary minerals present locally in variable amounts in the reydarfjordur borehole include calcite (3.4 w/m °c), chlorite (4–6 w/m °c), quartz (7.7 w/m °c), epidote (2.6–3.0 w/m °c) and haematite (about 11 w/m °c). in the lopra-1/1a borehole, zones of maximum temperature gradient and inferred minimum mean conductivity are generally found within the massive cores of the basalt flows characterised by high density and low porosity (fig. 5). this clearly points to a local increase in low-conductivity secondary minerals such as clay minerals (about 1.5– 2 w/m °c), analcite (1.3 w/m °c) and hydrous zeolite minerals like stilbite (1.2 w/m °c). in order to take a step further into the analysis of the temperature-gradient variability related to mineralogical variations, fig. 5 also presents a curve showing the relative intensity of red colour reflected from the formation. the colour information is extracted from a digital colour photograph with 24-bit resolution of a montage of cutting samples from the borehole (using the public domain program imagej). the relative intensity of red is computed from the values of the red, green and blue channels as the function red2/(green × blue). the main idea is to test without extensive mineralogical analysis whether the mineral haematite might play an important role. haematite (fe2o3) has a bright reddish colour and very high thermal conductivity. it is formed mainly by oxidation of primary magnetite and secondary iron hydroxides. the colour curve shows that many intervals of maximum reddish colour more or less coincide with intervals of low temperature gradient. a close correlation is seen particularly within the depth interval 1550–1700 m. an average cuttings lag time correction of 3 m is applied for the whole section. however, the lag varies with drilling rate, which varies with the hardness of the formation, and a locally better correlation may be obtained by applying a slightly different depth shift of the colour curve. from the continuously cored vestmanna-1 borehole, also in the faroe islands, the content of haematite in highly oxidised tuffaceous claystone may be as high as 25% estimated from bulk rock chemistry (unpublished data, r. waagstein). a unit of highly altered basalt or tuff consisting of silicate minerals like pyroxene, plagioclase, clay and zeolites with an assumed average matrix thermal conductivity of about 2 w/m °c plus 25% of haematite (conductivity about 11 w/m °c) will have a bulk conductivity of about 3 w/m °c , as calculated using the geometric mean formula, (see above). this is sufficiently high to produce the lowest temperature gradients of about 20°c/km. nonetheless, units of increased thermal conductivity generally also have a high (apparent) neutron porosity, which requires minerals of high hydrogen content. this means that, although haematite may play an important part, other components are contributing and further studies are needed for a better understanding of the relation between rock thermal properties and secondary mineralogical components. estimates of heat flow basalts have low concentrations of the heat producing isotopes u, th and k, resulting in low heat production, generally within the range of 0.2–0.6 × 10–6 w/m3 (e.g. verdoya et al. 1998; chiozzi et al. 2003). the contribution to surface heat flow from a 3.5 km deep section is thus very small, of the order of 1–2 mw/m2. if not significantly perturbed by effects of topography, groundwater flow or potential long-term palaeoclimatic surface-temperature variations, heat flow should be almost constant along the drilled section. for sections of the lopra-1/1a borehole where temperature measurements are assumed to represent conductive equilibrium values, heat flow may therefore be estimated from the product of mean temperature gradient and mean characteristic thermal conductivity. there seems to be no significant perturbing effects at depths below 1300 m. the lopra-1/1a borehole is in an geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19103 104 area of small topographic height variations and the effect of topography upon temperature and temperature gradient was modelled to be insignificant. temperature perturbations are below 1°c. the artesian flow of water in the borehole is localised to levels above 1150–1200 m. as discussed above, both observations and model calculations show that the influence of palaeoclimatic surface temperature variations is insignificant at depths greater than 1200– 1300 m. mean temperature gradients from long depth intervals vary within the narrow range of 28–33°c/km (table 3) and are thus well defined. the main source of error in estimating heat flow is thus the choice of mean characteristic thermal conductivity. the uncertainty arises from the presence of local variations of the temperature gradient interpreted in the previous section in terms of conductivity variations associated with mineralogical changes. these changes are difficult to quantify in detail and thus not fully understood. by transferring thermal conductivity from laboratory measurements to representative in situ values, temperature and pressure dependency needs consideration. however, for basalt this dependency is small compared to most other crystalline rocks. the decrease of thermal conductivity of basalt with increasing temperature is of the order of only 5–10% for a temperature increase from 20°c in the laboratory to a temperature of 50–100°c in a borehole (cf. compilations in kappelmeyer & haenel 1974). a conductivity decrease of this magnitude is likely to be almost compensated by an equivalent increase of conductivity with pressure. the slightly lower temperature gradient in the deeper parts of the borehole (table 3) may indicate a slight general increase in average thermal conductivity with depth. this increase may be explained by decreasing porosity resulting from secondary mineralisation. the most accurate large interval temperature gradient is the assumed conductive mean equilibrium gradient of 32.9°c/km between 1400 and 2175 m. using our rock thermal conductivity measurements in the range of 1.7– 1.9 w/m °c with a mean of about 1.8 w/m °c, we obtain a heat flow within the range of 56–63 mw/m2. for the deeper parts of the borehole between 2200 and 3430 m, the temperature gradient is between 30 and 31°c/km. the lithology is here represented by roughly equal amounts of basalt and lapilli-tuff. low-porosity lapilli-tuff may have a mean conductivity of about 2.0–2.2 w/m °c and basalt of about 1.8 to 1.9 w/m °c. this yields a mean conductivity of about 2.0 w/m °c and a heat-flow estimate close to 60 mw/m2. some local intervals of massive basalt at depths between 2000 and 2115 m have well-defined temperature gradients between 35 and 38°c/km with a mean value of 36°c/km. although massive basalt units are inferred to have a slightly reduced conductivity judged by their temperature gradients, average conductivity (inferred from our measurements) seems unlikely to be lower than 1.6 and not above 1.8 w/m °c, resulting in heat flow in the range 57–65 mw/m2. despite some uncertainty about details of the conductivity variations and their causes, we therefore estimate terrestrial heat flow for the lopra-1/1a borehole to be of the order 60 ± 5 mw/m2. this value is about 15 mw/m2 higher than a previous estimate of balling et al. (1984). the main reason is new measurements showing higher values of thermal conductivity and also the recognition that the neutron log data cannot be interpreted in terms of intervals of real high porosity and water-filled pores resulting in reduced thermal conductivity. on the contrary, local intervals of apparent high porosity are generally observed as having low temperature gradients and thus increased thermal conductivity. nearby areas of continental crust in the faroe–shetland basin south-east of the faroe islands have presentday heat flows between 45 and 65 mw/m2 (iliffe et al. 1999). in continental areas off the norwegian coast, mean heat flow is between 50 and 65 mw/m2 (sundvor et al. 2000). a heat flow value of around 60 mw/m2 is thus consistent with continental crust underlying the faroe islands. heat flow from the mantle in scandinavia is estimated to be around 25–35 mw/m2 (balling 1995). if similar values apply here, about 30 mw/m2 must originate from heat produced by decay of radiogenic isotopes in the crust, which requires a crust of continental composition. otherwise, a significant cooling component and/ or significantly increased mantle heat flow must be assumed, for which there is no other evidence. however, this must be the case for areas of oceanic crust north of the faroe islands, where heat flow generally between 60 and 75 mw/m2 is observed (sundvor et al. 2000). summary and conclusions the lopra-1/1a borehole drilled to a depth of 3.5 km offers a unique opportunity of obtaining accurate information on the thermal structure to a great depth in the faroe islands. high-precision temperature logging was carried out to a depth of 2175 m almost 13 years after drilling. temperatures in the upper 1200–1300 m are significantly disturbed by upward flow of ground water inside the borehole. for deeper levels, between 2175 m and total depth, only temperature logs from the commercial loggeus bulletin no 9 7 juli.pmd 07-07-2006, 14:19104 105 ging runs measured a relatively short time (up to 53 hours) after drilling fluid circulation are available. these temperatures have been corrected for the estimated effect of disturbances. the deepest point of accurate temperature information is 3430 m with a measured temperature of 98.6°c and a corrected, estimated equilibrium temperature of 103.5°c. temperature gradients calculated for depth intervals greater than 500–1000 m show only small variations between 28 and 33°c/km. the least-squares mean gradient for the undisturbed part of the borehole (1400–3430 m) is 31.4°c/km. levels of inflow of water to the upper part of the borehole are seen as major peaks on the 5 m mean interval temperature gradient. in addition, significant local temperature gradient variability is observed in the highprecision log between about 1250 m with minimum values down to 20–25°c/km and maximum values up to 45°c/ km. the latter variations correlate closely with variations in other logging parameters and inferred lithological variations within the lava succession and cannot be explained by ground-water flow. intervals of low temperature gradient generally match intervals of low density, high neutron porosity and increased borehole calliper and intervals of high temperature gradients match intervals of high density, low neutron porosity and decreased calliper. the observed correlation with neutron porosity is surprising. in a conductive regime, the temperature gradient should increase in lithological units of high porosity and pores filled with free water of low thermal conductivity. here, we observe that units of apparent high porosity have high gradients. this leads us to conclude that logged high neutron porosity does not represent real high porosity units with pores filled with free water. instead, materials of relatively high thermal conductivity compared to normal basaltic material must be present in significant amounts. the local temperature gradient variations are thus inferred to originate from variations in thermal conductivity. the latter variation is ascribed to secondary mineralisation and mineral alterations. this may produce both high conductivity minerals such as calcite, chlorite, quartz, epidote and haematite and low conductivity hydrous minerals such as clay and zeolite minerals. such inferred local variations in rock thermal properties are only partly reflected in our thermal conductivity measurements on core materials from the lopra-1/1a borehole and samples from surface outcrops in the lopra-1/1a area. these results, mostly on basalts and some on lapilli-tuff and tuff, show homogeneous conductivity with only small variations and mean values at about 1.8 (basalt) and 1.9 w/m °c (lapilli-tuff/tuff ). our measured lapilli-tuffs and tuffs generally show matrix (grain) conductivity about 20% higher than the basalts. the elevated conductivity of the former rocks may be explained by the abundance of secondary minerals with higher bulk conductivity than basalt. the increased conductivity in the originally porous part of flow units may be explained in a similar way by secondary mineralisation, as mentioned above. in some distinctly reddish intervals, haematite seems to contribute significantly to the increase in conductivity. however, further studies are needed in order to obtain a better understanding of the correlation between rock thermal properties and mineralogical alterations. because of the overall homogeneous mean temperature gradient structure, it is possible to obtain some information on palaeo-surface temperatures during the last glaciation by extrapolation of the temperature–depth function below 1200–1400 m to the ground surface. extrapolating the depth interval 1400–2175 m, a surface intercept of –6.7°c is obtained. this is 13–14°c below present-day surface temperature. although the extrapolation must be considered preliminary and approximate by nature, it suggests a temperature increase of the order of 12–16°c at the termination of the last glacial period. from well-defined temperature gradients and information on mean characteristic thermal conductivity of the drilled basaltic sequences, we estimate a conductive heat flow at the lopra-1/1a drill site of about 60 ± 5 mw/m2. this is about 15 mw/m2 higher than the previous estimate from the original borehole. the revised estimate is due mainly to new, higher thermal conductivity measurements and higher estimates of the conductivity of the porous parts of the basalt flows by taking secondary mineralisations into account. a heat flow value of about 60 mw/m2 is consistent with the faroe islands being underlain by continental crust. from our analysis we may conclude that the thermal regime and our reported temperatures, temperature gradients and heat-flow value from below a depth of 1200– 1400 m represent conductive equilibrium conditions without significant disturbances from the effect of drilling, ground-water flow or palaeoclimatic surface temperature variations. temperature structure, temperature gradients and heat flow may thus be taken as representative of a larger area around the drill site with similar basaltic lithology. with respect to heat flow, an assumption of similar lithology may not be necessary. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19105 106 acknowledgements valuable comments from two referees, andrea förster, gfz, potsdam and torben bidstrup, geus, copenhagen are gratefully acknowledged. this study has been supported by funds from the danish natural science research council. references balling, n. 1995: heat flow and thermal structure of the lithosphere across the baltic shield and northern tornquist zone. tectonophysics 244, 13–50. balling, n., kristiansen, j.i., breiner, n., poulsen, k.d., rasmussen, r. & saxov, s. 1981: geothermal measurements and subsurface temperature modelling in denmark. geoskrifter 16, 172 pp. århus, denmark: university of aarhus. balling, n., kristiansen, j.i. & saxov, s. 1984: geothermal measurements from the vestmanna-1 and lopra-1 boreholes. in: berthelsen, o., noe-nygaard, a. & rasmussen, j. (eds): the deep drilling project 1980–1981 in the faeroe islands. annales socientatis scientiarum faeroensis, supplementum ix, 137–148. tórshavn: føroya fróðskaparfelag. beck, a.e. 1988: methods for determining thermal conductivity and thermal diffusivity. in: haenel, r., rybach, l. & stegena, l. (eds): handbook of terrestrial heat-flow density determinations, 87–124. dordrecht: kluwer academic publishers. beck, a.e. & balling, n. 1988: determination of virgin rock temperatures. in: haenel, r., rybach, l. & stegena, l. (eds): handbook of terrestrial heat-flow density determinations, 59–85. dordrecht: kluwer academic publishers. berthelsen, o., noe-nygaard, a. & rasmussen, j. (eds) 1984: the deep drilling project 1980–81 in the faeroe islands. annales societas scientiarum færoensis, supplementum ix, 159 pp. tórshavn: føroya fróðskaparfelag. bott, m.h.p., sunderland, j., smith, p.j., casten, u. & saxov, s. 1974: evidence for continental crust beneath the faeroe islands. nature 248, 202–204. chiozzi, p., pasquale, v. & verdoya, m. 2003: heat from radioactive elements in young volcanics by γ-ray spectrometry. journal of volcanology and geothermal research 119, 205–214. dahl-jensen, d., mosegaard, k., gundestrup, n., clow, g.d., johnsen, s.j., hansen, a.w. & balling, n. 1998: past temperatures directly from the greenland ice sheet. science 282, 268– 271. gariépy, c., ludden, j. & brooks, c. 1983: isotopic and trace element constraints on the genesis of the faeroe lava pile. earth and planetary science letters 63, 257–272. hald, n. & waagstein, r. 1983: silicic basalts from the faeroe islands: evidence of crustal contamination. in: bott, m.h.p. et al. (eds): structure and development of the greenland–scotland ridge, 343–349. new york: plenum press. hald, n. & waagstein, r. 1984: lithology and chemistry of a 2 km sequence of lower tertiary tholeiitic lava drilled on suðuroy, faeroe islands (lopra-1). in: berthelsen, o., noe-nygaard, a. & rasmussen, j. (eds): the deep drilling project 1980–81 in the faeroe islands. annales societas scientiarum færoensis, supplementum ix, 15–38. tórshavn: føroya fróðskaparfelag. holm, p.m., hald, n. & waagstein, r. 2001: geochemical and pb-sr-nd isotopic evidence for separate hot depleted and iceland plume mantle sources for the palaeogene basalts of the faroe islands. chemical geology 178, 95–125. horai, k. 1971: thermal conductivity of rock-forming minerals. journal of geophysical research 76, 1278–1308. horai, k. 1991: thermal conductivity of hawaiian basalt: a new interpretation of robertson and peck’s data. journal of geophysical research 96, 4125–4132. huang, s., pollack, h.n. & shen, p.-y. 2000: temperature trends over the last five centuries reconstructed from borehole temperatures. nature 403, 756–758. iliffe, j.e., robertson, a.g., ward, g.h.f., wynn, c., pead, s.d.m. & cameron, n. 1999: the importance of fluid pressures and migration to the hydrocarbon prospectivity of the faeroe–shetland white zone. in: fleet, a.j. & boldy, s.a.r. (eds): petroleum geology of northwest europe: proceedings of the 5th conference, 601–611. london: geological society. jørgensen, o. 1984: zeolite zones in the basaltic lavas of the faeroe islands. in: berthelsen, o., noe-nygaard, a. & rasmussen, j. (eds): the deep drilling project 1980–81 in the faeroe islands. annales societas scientiarum færoensis, supplementum ix, 71– 91. tórshavn: føroya fróðskaparfelag. jørgensen, o. 2006: the regional distribution of zeolites in the basalts of the faroe islands and the significance of zeolites as palaeotemperature indicators. geological survey of denmark and greenland bulletin 9, 123–156 (this volume). kappelmeyer, o. & haenel, r. 1974: geothermics with special reference to application. geoexploration monographs. berlin: gebrüder borntraeger. kristiansen, j.i. 1991: nepr: a fortran-77 program for determining thermal conductivity and diffusivity by needle-probe inversion. computers & geosciences 17, 351–390. kukkonen, i.t. & joeleht, a. 2003: weichselian temperatures from geothermal heat flow data. journal of geophysical research 108, http://dx.doi.org/10.1029/2001jb001579. larsen, l.m., waagstein, r., pedersen, a.k. & storey, m. 1999: trans-atlantic correlation of the palaeogene volcanic successions in the faeroe islands and east greenland. journal of the geological society (london) 156, 1081–1095. oxburgh, e.r. & agrell, s.o. 1982: thermal conductivity and temperature structure of the reydarfjordur borehole. journal of geophysical research 87, 6423–6428. rasmussen, j. & noe-nygaard, a. 1970: geology of the faeroe islands. danmarks geologiske undersøgelse 1. række 25, 142 pp. richardson, k., smallwood, j.r., white, r.s., snyder, d.b. & maguire, p.k.h. 1998: crustal structure beneath the faroe islands and the faroe–iceland ridge. tectonophysics 300, 159– 180. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19106 107 robertson, e. & peck, d. 1974: thermal conductivity of vesicular basalt from hawaii. journal of geophysical research 79, 4875– 4888. roy, s. & rao, r.u.m. 1999: geothermal investigations in the 1993 latur earthquake area, deccan volcanic province, india. tectonophysics 306, 237–252. skogseid, j., planke, s., faleide, j.i., pedersen, t., eldholm, o. & neverdal, f. 2000: ne atlantic continental rifting and volcanic margin formation. in: nøttvedt, a. et al. (eds): dynamics of the norwegian margin. geological society special publication (london) 167, 295–326. sundvor, e., eldholm, o., gladczenko, t.p. & planke, s. 2000: norwegian–greenland sea thermal field. in: nøttvedt, a. et al. (eds): dynamics of the norwegian margin. geological society special publication (london) 167, 397–410. verdoya, m., pasquale, v., chiozzi, p. & kukkonen, i.t. 1998: radiogenic heat production in the variscan crust: new determinations and distribution models in corsica (northwest mediterranean). tectonophysics 291, 63–75. waagstein, r. 1988: structure, composition and age of the faeroe basalt plateau. in: morton, a.c. & parson, l.m. (eds): early tertiary volcanism and the opening of the ne atlantic. geological society special publication (london) 39, 225–238. manuscipt received 4 may 2005; revision accepted 9 september 2005. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19107 geological survey of denmark and greenland bulletin 7, 2004, p 77-80 77 small-scale mining is the main source of income for about 100 million people in asia, africa and south america. however, the processing of raw materials during this mining activity results in the release of large amounts of mercury to the environment, creating serious environmental problems. small-scale mining, or artisanal mining, is exploitation using only shovels, picks and hammers, carried out by individuals or small groups. a wide variety of commodities are exploited in this way, ranging from gold, diamonds, precious stones, tin, coal, dimension stones and slate. small-scale mining is often carried out by labourers with virtually no knowledge of safety procedures. tunnel cave-ins leading to loss of life are common, and the widespread use of mercury in gold extraction causes many long-term health problems for the miners. it is estimated that about 650 tonnes of mercury are annually released during small-scale mining to the environment, and this figure is likely to increase in the future. mercury is highly toxic and its use causes health problems not only for the miners, but also to the entire population in areas where small-scale mining takes place. some miners are aware of the dangers of using mercury, but have no knowledge of recycling procedures. several international organisations, such as the world bank, unido (united nations industrial development organisation), ilo (international labour organisation) and undp (united nations development programme), have launched programmes to examine the problems associated with small-scale mining. progress so far has been slow, and much more international awareness of the global mercury pollution of the environment from smallscale mining is required. the geological survey of denmark and greenland (geus) has worked as consultant to the world bank on projects involving small-scale mining in kyrgyzstan, mongolia and laos, and has also undertaken programmes concerning small-scale mining in lesotho for undp and in tanzania for the danish international development agency (danida). this paper reports on some of the initiatives carried out in kyrgyzstan and mongolia, to secure and sustain the small-scale mining industry in these regions. a historical view small-scale mining has been carried out since ancient times. many know the legend of jason and the golden fleece which may, in fact, be the first record of small-scale mining in the history of mankind. herdsmen traditionally placed a ram’s fleece in a river at springtime, which trapped gold particles tumbling down the river during the spring floods. after a month or so the fleece would be burned, or dried and shaken, to collect the gold grains. this ancient technique is still used by herdsmen in kyrgyzstan to recover gold from streams. small-scale mining – hazards and opportunities in kyrgyzstan and mongolia peter w.u. appel geological survey of denmark and greenland bulletin 7, 77–80 (2005) © geus, 2005 fig. 1. sluice for recovering placer gold in a stream in kyrgyzstan. the method of using a fleece has, however, largely been replaced by more efficient techniques. gold has a high specific gravity of 19 g/cm3, that makes gold easy to concentrate by mechanical means. the traditional gold pan, known from the gold rush in alaska in the late 1800s, is still one of the most efficient and widespread ways of recovering gold on a small scale. in a gold pan, light minerals such as quartz and feldspar are removed, and the heavy minerals, including grains of gold, are left in the pan. the larger gold grains can be hand-picked from the heavy minerals in the pan, but the fine-grained gold (gold dust) that can account for up to 60% of the gold may be lost. this is, of course, not very efficient, and the most widely used method to increase the recovery of gold is by amalgamation using mercury (see below). small-scale gold mining small-scale mining for gold may be carried out in situ on gold-bearing hard rock, or in gold placer deposits forming riverbeds and riverbanks. where tunnels are dug into the unconsolidated placer sediments, cave-ins are common and make working conditions very dangerous. the gravel that is extracted is sieved and the heavy minerals are then concentrated in a sluice (fig. 1). river water flowing through the sluice removes the light minerals, and large gold grains can then be hand-picked and the residue concentrated in a gold pan. this concentrate may again be hand-picked for visible gold grains, or treated with mercury (amalgamation). hard rock mining is commonly focussed on gold-bearing quartz veins. vertical shafts may be sunk down several tens of metres, and a network of tunnels up to hundreds of metres long branching out from the shaft are used to follow the goldbearing veins and excavate the ore. however, even here insufficient support of roofs and walls may result in frequent cave-ins and loss of life. the mined ore is crushed and ground by hand or in small mills. the fine material is then washed in a gold pan, hand-picked, and the residue often treated with mercury. amalgamation during amalgamation mercury is added to a mineral concentrate (fig. 2). gold dissolves in the mercury, forming an amalgam. the amalgam is then placed in an iron cup and heated over an open fire, which evaporates the mercury, and leaves behind the gold. amalgamation is a very efficient process for recovering fine-grained gold, but has the disadvantage that large amounts of mercury are released into the environment during the process. in mongolia alone several tonnes of mercury are released to the environment every year, and the amount released globally reaches many hundreds of tonnes. health problems mercury that evaporates during amalgamation condenses on cooling, and often finds its way into local streams and rivers as metallic mercury. metallic mercury may be converted by bacterial action into the even more toxic methylated mercury. both forms of mercury readily enter the food chain, and ultimately can lead to severe health problems for human beings. even small amounts of mercury affect the nervous system in human beings and cause tremors, while larger amounts can lead to erithism, a mental illness that leaves its victims irritated and very short-tempered. further stages in mercury poisoning give rise to inflammation in the gingiva, tunnel vision and permanent brain damage. mercury poisoning is particularly dangerous for pregnant women, as the foetus concentrates mercury selectively relative to the mother by a factor of up to ten. thus, while a mother may have only slight symptoms of mercury poisoning, the baby may be born with permanent brain damage. mercury levels of a population may be easily tested by analysing hair, blood or urine samples. recycling of mercury and alternative methods extracting gold by amalgamation has been used for centuries and on most continents. there are thus very large parts of the world which have become polluted, and consequently very many people that suffer from serious mercury poisoning. in south america various measures for recycling mercury have been experimented with, and alternative measures for extracting gold have been developed. 78 fig. 2. amalgamation in tanzania. note the shiny spots of metallic mercury in the pan. the most efficient, least expensive and most robust apparatus invented to recycle mercury is the so-called retort. this consists of segments of plumbing tubes joined together. the amalgam is placed in the retort cup on an open fire (fig. 3). the evaporating mercury escapes into the long thin tube that is cooled by a piece of wet cloth, and condenses back into mercury that is collected in a small glass filled with water. the condensed mercury can then be re-used. using retorts can reduce the release of mercury to the environment by more than 95%, and the method has been promoted by different organisations, including unido and the world bank. during world bank missions in kyrgyzstan and mongolia, the author held classes demonstrating how to use the retort (fig. 4), and distributed retorts to small-scale miners in mongolia. if mass-produced the cost of the retort is only 5 to 10 euros, which should allow funding organisations to hand them out to small-scale miners. teaching small-scale miners to use retorts is a major advance, but local regulations can prevent its implementation. during a world bank mission in kyrgyzstan in 2003, the author suggested that training courses for miners in the use of retorts should be established. the use of mercury is, however, forbidden by law in kyrgyzstan, and although mercury is actually used, teaching small-scale miners how to use retorts is forbidden. the mongolian government has a more relaxed attitude. use of mercury in mongolia is also forbidden, but teaching about the advantages of using retorts is not. there are also alternative methods of extracting gold from heavy mineral concentrates from placer deposits. a very simple method, called acid treatment, is used in kyrgyzstan. the heavy mineral concentrate is dried and treated with a hand magnet that removes all magnetic minerals, leaving only gold and metal sulphides. concentrated nitric acid is then added and the mixture is heated (fig. 5); the acid dissolves the sulphides such that only gold is left. unfortunately, when nitric acid dissolves the sulphides brown fumes of very toxic nitrogen oxides are given off. however, as long as the process is carried out in the open air and away from houses, the toxic gases pose little risk to people. the nitrogen oxides are very unstable and disintegrate rapidly into nitrogen and oxygen. sources of mercury mercury is a fairly rare metal, which only occurs in commercial quantities in a few places on earth. it has been mined for centuries in spain, algeria, china and kyrgyzstan. however, due to the globally declining demand for mercury, many 79 fig. 3. retort; the amalgam is placed in the cup that is heated over an open fire. evaporated mercury is condensed in the pipe and can be collected for recycling. match for scale. fig. 4. teaching small-scale miners in mongolia to use retorts. fig. 5. acid treatment. boiling the gold concentrate with concentrated nitric acid. the process releases very toxic brown fumes of nitrogen oxides. cup 80 mercury mines have either been closed down or production has decreased. nevertheless, there are apparently almost limitless supplies of cheap mercury available, as mercury has for more than a century been used as a catalyst in the production of chlorine and alkali metals in europe and north america. production methods that use mercury as a catalyst have recently been prohibited by the european union, and european factories have developed less toxic methods. although the ban on the use of mercury for production purposes does not come into full effect until 2020, many factories have already converted to new techniques. this has created a situation where approximately 50 000 tonnes of mercury in europe are surplus to requirements. such large quantities pose a major problem for the factories that hold the mercury, and they are actively seeking for ways of disposal. the current status of mercury trade within europe and the trade with mercury between europe and the rest of the world is described by maxon (2004) in a report for the european commission. the mercury trade within the european union (eu) is illustrated in fig. 6 which shows that export of mercury from the eu is mainly to third world countries in latin america, the caribbean, south asia and the pacific. it is evident that dutch, english and spanish companies in particular are helping european factories to dispose of their mercury stocks. what can be done to help small-scale miners? the most efficient way to reduce the health and environmental problems for miners is to teach them either to recycle mercury by using retorts, or to use alternative methods. with that purpose in mind, short teaching and training programmes for small-scale miners in bornuur township in mongolia have been carried out showing how to use retorts (fig. 4). this town is especially appropriate as a target for teaching, since a recent survey has shown that the population in bornuur has high contents of mercury in urine samples and that the mercury stems from the use of amalgamation by small-scale miners (tumenbayar 2003). building on these experiences, further projects are now being considered, including teaching and training medical doctors to diagnose symptoms of mercury poisoning. included in these programmes would also be the establishment of local information centres where small-scale miners and others can seek information on mercury problems. the global threat for the environment and health posed by the use of mercury can be reduced considerably and at little cost by teaching and training of miners, and handing out retorts that cost only a few euros. references maxon, p. 2004: mercury flows in europe and the world: the impact of decommissioned chlor-alkali plants. brussels: european commission. directorate general for environment, 104 pp. tumenbayar, b. 2003: action research on mercury pollution in boroo area, mongolia. ulaan baatar, mongolia: japan international cooperation agency, mongolia office, 75 pp. author’s address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: pa@geus.dk eu 10528 0 83 16860 108 16 48 429 7 403 774 243 245 the netherlands belgium germanyuk finland spain fig. 6. mercury trade in the european union (metric tonnes) in 2000. red arrows indicate export to mainly third world countries from the uk, spain, the netherlands and germany. geological survey of denmark and greenland bulletin 28, 2013, 65-68 65 a new seamless digital 1:500 000 scale geological map of greenland mikael pedersen, willy l. weng, nynke keulen and thomas f. kokfelt for around 40 years, the geological survey of greenland (ggu) and later the geological survey of denmark and greenland (geus) conducted systematic geological overview mapping of greenland, leading to the production of 14 printed map sheets on a scale of 1:500 000 which covers the entire country (fig. 1). the mapping was completed in 1999 and the map sheets were published between 1971 and 2004. revised 2nd editions of two of the sheets have been published (table 1). as a result of the development of the internet, the publication platform for geological maps has increasingly moved from traditional paper sheets to digital publications for most geological surveys worldwide. this tendency has changed the requirements to the underlying production, storage and publication platforms for map data. for the same reason, it has over the past decade been a strategy of geus to bring the 1:500  000 geological maps of greenland together as a seamless, digital product with a homogeneous legend. this has been further motivated by the increasing number of geographic information system (gis) and web applications to which such a map would provide a valuable background, as well as to many international projects such as onegeology. the harmonisation of the geological maps was carried out by a project team at geus comprising both geological and gis expertise. the first version of the new seamless map was finished in 2012 and will be released in various web applications in 2013. digitisation the production of a seamless, digital product makes new demands on the underlying map data. most importantly, all data have to be in vector form in a gis environment. only the three most recent of the 14 original map sheets mentioned above were produced digitally (sheets 6, 9, and 11; fig. 1, table 1). as part of the project thematic maps and data of north and northeast greenland: geology, mineral occurrences and hydrocarbons (mikkelsen et al. 2005; christoffersen & jepsen 2007) four of the printed map sheets (sheets 7, 8, 10 and 12) were vectorised. of the remaining seven sheets, two sheets (1 and 3) were up for revision, and new editions were produced digitally in 2007 and 2010. the last five (sheets 2, 4, 5, 13 and 14) were vectorised in order to create full digital coverage. the vectorisation was done from scanned versions of the published maps. in this process, the geological information was taken from the maps and integrated with the best available topographic base-map data, which in most cases were photogrammetric data derived from 1:150 000 aerial photographs (based on aerotriangulations of varying quality). some of the topography from printed maps was taken from g/250 vektor data from the danish geodata agency and used in map sheets 1, 4 and 5. due to the nature of the mapping technique and the geodetic control, the old geological paper maps had a much better local than absolute accuracy: the geometric relations between a geological feature and 6 7 9 10 11 12 13 14 500 km1 2 3 4 5 8 fig. 1. index map of greenland showing the boundaries of the 14 geological map sheets on a scale of 1:500 000 covering greenland. © 2013 geus. geological survey of denmark and greenland bulletin 28, 65–68. open access: www.geus.dk/publications/bull 6666 the nearby topographic features, such as shore lines and rivers, are usually correct even if the location of such features is quite wrong (as indicated for example by the geographic grid). during vectorisation in the gis environment, it was therefore constantly necessary to shift the scanned geological map in order to fit the printed map topography to the new digital topography, which in some cases was rather different from that on the printed map. the digitisation included boundaries between exposed geological units and structural elements, all of which were attributed with type and source information. subsequent to the vectorisation, polygons were formed and encoded according to their geological type. as the gis data were stored in a single seamless database, geometric ambiguities across map-sheet boundaries had to be corrected. the geometric quality of the final map varies substantially. in areas with well-defined topographic features and the best possible photogrammetry, mean errors around 50 m can be expected. however, in areas where topography from paper maps was used as reference during digitisation, errors fig. 2. geological map of the mestersvig area in central east greenland. a, c: examples of how mapping inconsistencies between two printed map sheets are displayed, where the same unit is shown with different colours across the former map-sheet boundary. b: an example of a geometrical problem where a boundary between two units shows a ‘jump’. these small inconsistencies remain on the new digital map. no name compilers published edition cartographic topography technique source 1 sydgrønland a.a. garde 2007 2 digital mixed 2 frederikshåb isblink – søndre strømfjord j.h. allaart 1982 analog photogr. 3 søndre strømfjord – nuussuaq a.a. garde & m. marker 2010 2 digital photogr. 4 upernavik isfjord j.c. escher 1985 analog paper map 5 thule p.r. dawes 1991 analog paper map 6 humboldt gletscher p.r. dawes & a.a. garde 2004 digital photogr. 7 nyeboe land n. henriksen 1989 analog (d) photogr. 8 peary land h.-j. bengaard & n. henriksen 1986 analog (d) photogr. 9 lambert land h.f. jepsen 2000 digital photogr. 10 dove bugt n. henriksen 1997 analog (d) photogr. 11 kong oscar fjord j.c. escher 2001 digital photogr. 12 scoresby sund h.-j. bengaard & n. henriksen 1984 analog (d) photogr. 13 kangerdlugssuaq j.s. myers, p.r. dawes & t.f.d. nielsen 1988 analog photogr. 14 skjoldungen j.c. escher 1990 analog photogr. table 1. data on the greenlandic bedrock maps 1:500 000 d: vectorised as a part of the project thematic maps and data of north and northeast greenland (mikkelsen et al. 2005; christoffersen & jepsen 2007) greenland a b 10 km c 67 are probably around 250 m. the position of the ice margin shown on the paper maps is even worse. in the baffin bay region for example, the ice-margin position in the 1950s is used. another important part of the geographic data set is the place names. although not all of the place names on the printed maps have been digitised yet, the current data set already contains more than 2500 place names. the concept of map scale also remains when data are being used digitally. the target 1:500 000 scale has governed the selection of features and their degree of detail; it also applies to the base topography that was duly generalised, typically from 1:100 000 originals. geological harmonisation the 14 map sheets that form the basis of the new seamless 1:500 000 scale map were published between 1982 and 2010. the original maps all have their own way of presenting the geology, depending on the map compilers and the individual mapping geologists. the maps vary greatly in degree of detail reflecting the amount of time spent during the original mapping, and the availability of helicopter support in the different areas. furthermore, the maps show an evolution in the understanding of the geology, and changes in mapping philosophy from older to younger maps. an important issue was to implement a uniform geological nomenclature to the entire map, so that a single, harmonised legend for all of greenland could be developed. since no geological remapping was done, the level of detail displayed in different areas remains uneven. therefore, boundaries between mapping areas of individual mapping parties, both within former map areas and across former map boundaries, to a large extent remain visible. an example of this is shown in fig. 2. the homogenised legend for the thematic maps of north and northeast greenland at 1:250  000 (christoffersen & jepsen 2007) was used used as a basis for further harmonisation; the harmonisation process was started by combining the adjacent four map sheets (sheet 5, 6, 12 and 13), where similar geological units occur, into a common legend. the north-western, northern, north-eastern and eastern parts of greenland largely consist of palaeoproterozoic to paleogene sedimentary basins. the sediments are divided into groups, formations and members described in variable, although generally high degrees of detail. the western, southern and south-eastern parts of greenland mainly consist of metamorphosed archaean and palaeoproterozoic fig. 3. the colours on the digital map follow the colour scheme of the original 1:500 000 scale maps, but have been updated to modern standards where necessary. an example from south-east greenland (map sheet 14). a: the new digital map. b: the same area on the printed version that was published in 1990 (table 1). a b 10 km greenland 6868 basement rocks. in line with the original, printed 1:500 000 maps, the rock units in these parts of greenland are mainly divided by their age and petrography. in cases where archaean rocks have undergone later reworking, e.g. archaean gneisses deformed and metamorphosed in the palaeoproterozoic, the entry in the legend is placed under the time of formation and not under the time of reworking, as on the original paper map sheets. the legends of most of the printed map sheets contain a number of rock types with generic names such as orthogneiss, amphibolite, mica schist, granite and ultramafic rocks. these units have been pooled into broader units that cover all of the corresponding rock types of similar age. however, if the original map sheet showed similar rock units with special features as separate types (e.g. the welldescribed, pyroxene-bearing ilivertalik granite), these have been retained in order not to loose information compared with the original map sheets. legend the link between geological features on a digital geological map and the corresponding legend is ensured by the use of codes. each polygon on the map has a code, which is translated into a colour and a descriptive text by the gis program. in the course of geological harmonisation, the codes from the various input maps were translated into a new, homogeneous encoding scheme so that, e.g. an ‘undifferentiated gneiss’ in one area got the same code as a corresponding unit in another area. this work was carried out using a combination of translation tables and python scripts, and resulted in a total list of 443 unique codes each representing a geological unit. the geological units were subsequently compiled in a legend organised by age and region. rock units that belong to a supergroup, a sedimentary basin, an igneous province or a metamorphic complex have been grouped together, even if they cover more than one time period. the grouping by age starts with quaternary deposits and ends with eoarchaean rocks of the isua complex. each original 1:500 000 map sheet legend had its own style, therefore a harmonisation of rock descriptions was also necessary. where the nomenclatures are obsolete, rock names have been adopted to modern nomenclature. we mainly followed the principles by the british geological survey (gillespie & styles 1999; hallsworth & knox 1999; robertson 1999). a data set containing as many different feature types as described requires careful symbolisation. the one currently used tries to stay as close as possible to the tradition of the printed maps of ggu and geus. however, some colours – mainly for igneous rocks – have been adapted to colours that are more common on modern maps (fig. 3). the new digital compilation is to a large degree intended for viewing on a computer screen. for that reason, the first version of the new map has only been assigned rgb colours. another colour scheme for map-sheet printing using cmyk colours will be developed at a later stage. publication platforms compared to a printed map, a digital geological map has more potential applications. first of all, it is well suited for gis work where users can zoom in on areas of interest and examine the geology e.g. by clicking on geological features and receive detailed information from the underlying databases. the new seamless map has been integrated in such a web-gis application which can be found at http://data. geus.dk/map2/geogreen. the seamless map is furthermore an important geus contribution to the global onegeology portal (http://portal.onegeology.org). this portal aims to put geological maps from all countries in the world together on a scale of 1:1 million or better, by using distributed web map services (wms). the geological map of greenland on this portal is currently a 1:2 500 000 scale map. the new 1:500 000 map will fulfil geus’ participation in this important international initiative. references christoffersen, m. & jepsen, h.f. 2007: geological maps of north and north-east greenland 1:250 000. copenhagen: geological survey of denmark and greenland. gillespie, m.r. & styles, m.t. 1999: bgs rock classification scheme 1. classification of igneous rocks. british geological survey research report rr 99-06, 52 pp. hallsworth, c.r. & knox, r.w.o’b. 1999: bgs rock classification scheme. classification of sediments and sedimentary rocks. british geological survey research report rr 99-03, 44 pp. mikkelsen, n., jepsen, h.f., ineson, j.r., piasecki, s., von platen-hallermund, f., schøjth, f., thomassen, b. & weng, w.l. 2005: thematic maps and data of north and northeast greenland: geology, mineral occurrences and hydrocarbons. danmarks og grønlands geologiske undersøgelse rapport 2005/28, 56 pp. robertson, s. 1999: bgs rock classification scheme 2. classification of metamorphic rocks. british geological survey research report rr 99-02, 24 pp. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark; e-mail: mp@geus.dk geological survey of denmark and greenland bulletin 3, 172-196 172 concluding remarks in the last section on the environmental changes within the seven regions through the late quaternary, it has been demonstrated how the skagen sequence ‘moved’ into the present-day faunal community known from this area, ‘coming’ from older deposits in many ways, according to the molluscs, different from what hitherto was known in other parts of the danish area during the late quaternary. the development is graphically shown in fig. 93 (fold-out, back cover). regarding the climatic changes, recalling figs 102 and 103, the molluscs have given a clear record as far as the main trends are concerned – the interglacial– glacial cycle. however, the climatic changes during the holocene, if they were ever more than small, were overshadowed by the facies changes affecting the danish area. from the eemian as well, it must be concluded that differences in facies made the difference between the regions, and that the well-established more temperate eemian marine fauna was connected only with the shallow-water environment. among the 140 species recorded from the eemian, 118 or 84.3% occur in the holocene subfossil material and/or recent fauna. however, it should be noticed that among the 22 species only found in the eemian, ten species or 7.1% are purely lusitanian forms, which include the no longer extant species paphia senescens. the lusitanian forms are: one gastropod – haminoea navicula; one scaphopod – dentalium vulgare; and seven bivalves – mytilaster lineatus, mytilaster solidus, lucinella divaricata, plagiocardium papillosum, gastrana fragilis, abra segmentum and gouldia minima. along with the high percentage of purely lusitanian forms – compared with and not found in the holocene – it is shown that among the 90 boreo-lusitanian species from the eemian, by far the dominating group covering 64.3% of the fauna, 87 species occur also in the holocene. the weichselian marine fauna known from the older yoldia and younger yoldia sea deposits also has a characteristic of its own, with about one third of the fauna restricted to the yoldia seas. nearly half of them are either purely arctic, such as the two bivalves portlandia arctica and macoma torelli, or with arctic– subarctic affinities, such as the three gastropods alvania cruenta, turritella erosa and cylichna occulta, and the four bivalves bathyarca glacialis, macoma loveni, pandora glacialis, and lyonsia arenosa. one third of the weichselian fauna is found both among the eemian and the holocene species (subfossil and/or recent). for the last third, the majority (11 out of 17) are also recorded only from the holocene. when the weichselian marine fauna itself is looked at, comparing the older and younger yoldia sea faunas, 23 out of the 54 species are common to both, while one third is only found in the younger yoldia sea deposits from where the subarctic–boreal species among the bivalves are: nuculana minuta, arctica islandica, and zirfaea crispata. the subarctic–boreal– lusitanian species count one polyplacophor: tonicella marmorea; one gastropod: buccinum undatum; and two bivalves: mytilus edulis and macoma balthica. this demonstrates in the best way the boreo-arctic impact around 13 000 b.p. (14c years) in the shallow-water environment characterised by zirfaea crispata and mytilus edulis – the former giving name to the deposits of that time in vendsyssel. however, the main result of this investigation was the comparison between the fossil faunas and the molluscan faunas now living before our eyes – as c.g.j. petersen expressed it in 1910 – using the c.g.j. petersen bottom community concept step by step in the seven stages from the eemian to the subatlantic within the seven regions in the danish realm. acknowledgements this study was supported by a one-year grant from the carlsberg foundation. during that year – and the following years – the geological survey of denmark and greenland (geus) has supplied me with all the facilities needed for the research. the support of both these institutions is greatly appreciated. among the many helpful colleagues at the survey, i would like to thank lasse gudmundsson for keeping order in the many samples that were analysed, and frants von platen-hallermund for making the compilations seen in the figures and appendix listing the molluscan species and other data. state geologist at that time johnny fredericia, who caused me to take up the challenge and continued to support me is thanked, as is richard bradshaw, state geologist of my new department at the survey, the department of environmental history and climate geus bulletin no 3.pmd 28-06-2004, 08:46172 173 change. j. heinemeier contributed with appendix 4 on the 14c dates on shell macrofossils from the skagen cores. kaare l. rasmussen read an early draft of the work and the referees h.g. petersen and s. funder contributed to make the work better – all are thanked. with great experience in writing for me, birgit jørgensen did the typing and commented upon the english. susanne veng christensen has made the final copy. peter john crabb revised the english in the most thorough way. to the former curator of the vertebrate collection at the geological museum, university of copenhagen, now happily at the gram museum, ella hoch, this book is dedicated. references andersen, b.g. 1965: the quaternary of norway. in: rankama, k. 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ovata acanthocardia echinata (linnaeus 1758) cardium echinatum acar nodulosa (müller 1766) aclis ascaris (turton 1819) aclis minor (brown 1827) aclis walleri jeffreys 1867 acmaea tessulata (müller 1776) acmaea virginea (müller 1776) tectura virginea acteon tornatilis (linnaeus 1758) adipicola simpsoni (marshall 1900) admete viridula (fabricius 1780) aequipecten opercularis (linnaeus 1758) pecten opercularis akera bullata müller 1776 alvania abyssicola (forbes 1850) rissoa abyssicola alvania cimicoides (forbes 1844) rissoa cimicoides alvania cruenta odhner 1915 alvania jan mayeni (friele 1886) rissoa jan mayeni alvania jeffreysi (waller 1864) alvania lactea (michaud 1830) rissoa lactea alvania punctura (montagu 1803) rissoa punctura alvania scrobiculata (möller 1842) rissoa scrobiculata amauropsis islandicus (gmelin 1791) angulus fabulus → fabulina fabula angulus tenuis (da costa 1778) tellina tenuis anomia aculeata → heteranomia squamula anomia ephippium linnaeus 1758 anomia patelliformis → pododesmus patelliformis anomia squamula → heteranomia squamula antalis agile g.o. sars 1878 antalis entalis (linnaeus 1758) dentalium entalis aporrhais pespelicani (linnaeus 1758) aporrhais serresianus (michaud 1828) arca glacialis → bathyarca glacialis arcinella plicata → saxicavella jeffreysi arcopagia crassa (pennant 1778) arctica islandica (linnaeus 1767) cyprina islandica geus bulletin no 3.pmd 28-06-2004, 08:46180 181 assiminea grayana fleming 1828 astarte banksie → tridonta montagui astarte borealis → tridonta borealis astarte compressa → tridonta elliptica astarte sulcata (da costa 1778) axinopsida orbiculata (g.o.sars 1878) axinopsis orbiculata axinopsis orbiculata → axinopsida orbiculata axinus ferruginosus → leptaxinus ferruginosus axinus flexuosus → thyasira flexuosa balcis devians → vitreolina philippii barleeia unifasciata (montagu 1803) barnea candida (linnaeus 1758) pholas candida bathyarca glacialis (gray 1824) arca glacialis bathyarca pectunculoides (scacchi 1834) bela exarata g.o.sars 1818 bela incisula → oenopota incisula bela nobilis → oenopota turricola bela trevelliana → oenopota trevelliana bela turricola → oenopota turricola bela violacea → oenopota violacea bittium reticulatum (da costa 1778) cerithium reticulatum boreotrophon clathratus (linnaeus 1767) trophon clathratus boreotrophon truncatus (ström 1768) brachystomia carozzai van aartsen 1987 brachystomia eulimoides hanley 1844 odostomia eulimoides odostomia pallida buccinum cyaneum bruguière 1792 buccinum groenlandicum buccinum groenlandicum → buccinum cyaneum buccinum undatum linnaeus 1758 cadulus jeffreysi → cadulus subfusiforme cadulus subfusiforme (m. sars 1865) cadulus jeffreysi caecum glabrum (montagu 1803) calliostoma formosa (mighels 1842) calliostoma zizyphinum (linnaeus 1758) callochiton septemvalvis (montagu 1803) capulus ungaricus (linnaeus 1758) cardium ciliatum → clinocardium ciliatum cardium echinatum → acanthocardia echinata cardium edule → cerastoderma edule cardium edule → cerastoderma glaucum cardium exiguum → parvicardium exiguum cardium fasciatum → parvicardium ovale cardium groenlandicum → serripes groenlandicus cardium minimum → parvicardium minimum cardium nodosum → parvicardium scabrum cardium norvegicum → laevicardium crassum cardium papillosum → plagiocardium papillosum cerastoderma edule (linnaeus 1758) cardium edule cerastoderma glaucum (poiret 1789) cardium edule var. balticum cerithiella metula (lovén 1846) cerithiopsis barleei jeffreys 1867 cerithiopsis tubercularis (montagu 1803) cerithium reticulatum → bittium reticulatum chamelea striatula (da costa 1778) venus gallina chemnitzia lactea → turbonilla lactea chlamys islandica (o.f.müller 1776) pecten islandicus chlamys varia (linnaeus 1758) pecten varius chrysallida eximia (jeffreys 1849) parthenina eximia chrysallida decussata (montagu 1803) chrysallida indistincta (montagu 1808) parthenia indistincta chrysallida obtusa (brown 1827) parthenia interstincta chrysallida spiralis (montagu 1803) parthenia spiralis cingula proxima → onoba proxima cingula semistriata (montagu 1808) putilla semistriata rissoa semistriata cingula striata → onoba semicostata cingula turgida (jeffreys 1870) cingula vitrea → onoba vitrea circe minima → gouldia minima claturella linearis → raphitoma linearis clausinella fasciata (da costa 1778) venus fasciata clinocardium ciliatum (fabricius 1780) cardium ciliatum clione limacina (phipps 1774) cochlodesma praetenue (pulteney 1799) colus gracilis (da costa 1778) colus jeffreysianus (fischer 1868) colus sabini (gray 1824) corbula gibba (olivi 1792) crenella decussata (montagu 1803) crepidula fornicata (linnaeus 1758) cultellus pellucidus → phaxas pellucidus cupidaria cuspidata (olivi 1792) geus bulletin no 3.pmd 28-06-2004, 08:46181 182 cyamium minutum → turtonia minuta cylichna alba (brown 1827) cylichna cylindracea (pennant 1777) cylichna occulta (mighels 1841) cylichna scalpta cylichna propinqua cylichna propinqua → cylichna occulta cylichna scalpta → cylichna occulta cypraea europaea → trivia monacha cyprina islandica → arctica islandica cytharella coarctata (forbes 1840) mangelia costata delectopecten vitreus (gmelin 1791) dentalium entalis → antalis entalis dentalium vulgare da costa 1778 devonia perrieri (malard 1904) diaphana hyalina → diaphana minuta diaphana minuta brown 1827 diaphana hyalina divaricella divaricata → lucinella divaricata donax vittatus (da costa 1778) dosinia exoleta (linnaeus 1758) dosinia lincta (montagu 1803) dosinia lupinus dosinia lupinus → dosinia lincta ebala nitidissima (montagu 1803) eulimella nitidissima emarginula fissura (linnaeus 1758) ensis arcuatus (jeffreys 1865) ensis ensis (linnaeus 1758) solen ensis ensis siliqua (linnaeus 1758) entalina tetragona (brocchi 1814) enteroxenos oestergreni bonnevie 1902 epitonium clathratulum (kanmacher 1797) epitonium clathrus (linnaeus 1758) scalaria communis epitonium trevelyanum (johnston 1841) epitonium turtonis (turton 1819) scalaria turtonae eulima bilineata (alder 1848) eulima distorta → vitreolina philippii eulimella acicula → eulimella laevis eulimella laevis (brown 1827) syrnola laevis eulimella acicula eulimella nitidissima → ebala nitidissima eulimella scillae (scacchi 1835) evalea divisa → ondina divisa fabulina fabula (gmelin 1791) angulus fabulus tellina fabula fusus antiquus → neptunea antiqua gari depressa (pennant 1777) psammobia vespertina gari fervensis (gmelin 1791) psammobia faeroeensis gari tellinella (lamarck 1818) gastrana fragilis (linnaeus 1758) gibbula cineraria (linnaeus 1758) trochus cineraria gibbula tumida (montagu 1803) trochus tumida glossus humanus (linnaeus 1758) gouldia minima (montagu 1803) circe minima graphis albida (kanmacher 1798) haliella stenostoma (jeffreys 1858) haminea navicula → haminoea navicula haminoea navicula (da costa 1778) haminea navicula hanleya hanleyi (bean 1844) helcion pellucidum (linnaeus 1758) nacella pellucidum patina pellucida hemiaclis ventrosa (jeffreys ms fricle 1874) heteranomia squamula (linnaeus 1758) anomia squamula anomia aculeata hiatella arctica (linnaeus 1758) saxicava arctica hiatella rugosa (linnaeus 1758) saxicava rugosa hinia incrassata (ström 1768) nassa incrassata hinia pygmaea (lamarck 1822) nassa pygmaea hinia reticulata (linnaeus 1758) nassa reticulata omalogyra atomus → omalogyra atomus hydrobia neglecta muus 1963 hydrobia stagnorum → hydrobia ventrosa hydrobia ulvae (pennant 1777) peringia ulvae hydrobia ventrosa (montagu 1803) hydrobia stagnorum iothia fulva (müller 1776) ischnochiton albus (linnaeus 1767) jujubinus clelandi (w. wood 1828) kellia suborbicularis (montagu 1803) kelliella miliaris (philippi 1844) kennerleya glacialis → pandora glacialis geus bulletin no 3.pmd 28-06-2004, 08:46182 183 lacuna crassior (montagu 1803) lacuna divaricata → lacuna vincta lacuna pallidula (da costa 1778) stenotis palidula lacuna parva (montagu 1803) lacuna puteolus lacuna puteolus → lacuna parva lacuna vincta (montagu 1803) lacuna divaricata laevicardium crassum (gmelin 1791) cardium norvegicum lamellaria perspicua (linnaeus 1758) leda minuta → nuculana minuta leda pernula → nuculana pernula lepeta caeca (müller 1776) lepidochitona cinereus (linnaeus 1767) leptaxinus ferruginosus (forbes 1844) axinus ferruginosus leptochiton asellus (gmelin 1791) lepton nitidum (turton 1822) lepton squamosum (montagu 1803) limacina balea → limacina retroversa limacina retroversa (fleming 1823) spiralis retroversus limacina balea spiralis balea limaria hians (gmelin 1791) limaria loscombi (sowerby 1832) limatula subauriculata (montagu 1808) liomesus ovum (turton 1825) liostomia clavula (lovén 1846) litorina rudis → littorina saxatilis litorina rudis → littorina tenebrosa littorina littorea (linnaeus 1758) littorina mariae sacchi & rastelli 1966 littorina obtusata (linnaeus 1758) littorina saxatilis (olivi 1792) litorina rudis littorina tenebrosa (montagu 1803) litorina rudis lucina borealis → lucinoma borealis lucina divaricata → lucinella divaricata lucinella divaricata (linnaeus 1758) lucina divaricata divaricella divaricata lucinoma borealis (linnaeus 1758) lucina borealis lucinopsis undata → mysia undata lunatia alderi (forbes 1838) natica intermedia natica alderi lunatia catena (da costa 1778) natica catena lunatia montagui (forbes 1838) lunatia pallida (broderip & sowerby 1829) natica groenlandica lutraria elliptica → lutraria lutraria lutraria lutraria (linnaeus 1758) lutraria elliptica lymnaea peregra (müller 1774) lyonsia arenosa (möller 1842) lyonsia norvegica (gmelin 1791) macoma balthica (linnaeus 1758) tellina balthica macoma calcarea (gmelin 1791) tellina calcaria macoma loveni jensen 1904 tellina loveni macoma torelli jensen 1904 tellina crassula mactra solida → spisula solida mactra stultorum (linnaeus 1758) mactra corallina cinerea mactra subtruncata → spisula subtruncata mactra corallina cinerea → mactra stultorum malletia obtusa (g.o. sars 1872) mangelia attenuata (montagu 1803) mangelia brachystoma (philippi 1844) mangelia costata → cytharella coarctata mangelia nebula (montagu 1803) margarites helicinus (phipps 1774) melanella alba (da costa 1778) melanella lubrica (monterosato 1891) melaraphe neritoides (linnaeus 1758) menestho divisa → ondina divisa modiola modiolus → modiolus modiolus modiolaria discors → musculus discors modiolaria laevigata → musculus laevigatus modiolaria marmorata → modiolaria tumida modiolaria nigra → musculus niger modiolaria tumida (hanley 1843) musculus tumidus modiolaria marmorata modiolula phaseolina (philippi 1844) mytilus phaseolinus modiolus adriaticus (lamarck 1819) mytilus adriaticus modiolus modiolus (linnaeus 1758) mytilus umbilicatus modiola modiolus montacuta bidentata → mysella bidentata montacuta ferruginosa → tellimya ferruginosa geus bulletin no 3.pmd 28-06-2004, 08:46183 184 montacuta substriata (montagu 1803) musculus discors (linnaeus 1767) modiolaria discors musculus laevigatus (gray 1824) modiolaria laevigata musculus niger (gray 1824) modiolaria nigra musculus tumidus → modiolaria tumida mya arenaria linnaeus 1758 mya truncata linnaeus 1758 myrtea spinifera (montagu 1803) mysella bidentata (montagu 1803) montacuta bidentata mysella dawsoni (jeffreys 1864) mysella tumidula (jeffreys 1867) mysia undata (pennant 1777) lucinopsis undata mytilaster lineatus (gmelin 1791) mytilaster solidus (poli 1795) mytilus minimus mytilus adriaticus → modiolus adriaticus mytilus edulis linnaeus 1758 mytilus minimus → mytilaster solidus mytilus phaseolinus → modiolula phaseolina mytilus umbilicatus → modiolus modiolus nacella pellucidum → helcion pellucidum nassa incrassata → hinia incrassata nassa pygmaea → hinia pygmaea nassa reticulata → hinia reticulata natica affinis (gmelin 1790) natica clausa natica alderi → lunatia alderi natica catena → lunatia catena natica clausa → natica affinis natica groenlandica → lunatia pallida natica intermedia → lunatia alderi neptunea antiqua (linnaeus 1758) fusus antiquus neptunea despecta (linnaeus 1758) neritina fluviatilis → theodoxus fluviatilis nototeredo norvegica (spengler 1792) nucella lapillus (linnaeus 1758) purpura lapillus nucula nitida → nucula nitidosa nucula nitidosa winckworth 1930 nucula nitida nucula nucleus (linnaeus 1767) nucula sulcata (bronn 1831) nucula tenuis → nuculoma tenuis nuculana minuta (müller 1776) leda minuta nuculana pernula (müller 1776) leda pernula nuculoma hanleyi winckworth 1931 nuculoma tenuis (montagu 1808) nucula tenuis obtusella alderi (jeffreys 1858) ocenebra erinacea (linnaeus 1758) odostomia acuta jeffreys 1848 odostomia albella lovén 1846 odostomia unidentata odostomia conoidea winckworth 1932 odostomia eulimoides → brachystomia eulimoides odostomia insculpta → ondina divisa odostomia pallida → brachystomia eulimoides odostomia plicata (montagu 1803) odostomia rissoides → odostomia scalaris odostomia scalaris macgillivray 1843 odostomia rissoides odostomia turrita hanley 1844 odostomia umbilicaris (malm 1863) odostomia unidentata → odostomia albella oenopota incisula (verrill 1882) bela incisula oenopota trevelliana (turton 1834) bela trevelliana oenopota turricola (montagu 1803) bela turricola bela nobilis oenopota violacea (mighels & adams 1842) bela violacea omalogyra atomus (phillippi 1841) omalogyra atomus ondina diaphana (jeffreys 1848) ondina divisa (j. adams 1797) menestho divisa evalea divisa odostomia insculpta onoba aculeus (gould 1841) onoba proxima (forbes & hanley 1850) cingula proxima onoba semicostata (montagu 1803) rissoa striata cingula striata onoba vitrea (montagu 1803) cingula vitrea rissoa vitrea ostrea edulis linnaeus 1758 ovatella myosotis (draparnaud 1801) palliolum greenlandicum (sowerby 1842) pecten groenlandicus palliolum striatum (müller 1776) geus bulletin no 3.pmd 28-06-2004, 08:46184 185 palliolum tigerinum (müller 1776) pandora glacialis leach 1819 kennerleya glacialis panomya arctica (lamarck 1818) paphia aurea (gmelin 1791) tapes aureus paphia aurea senescens (cocconi 1873) tapes senescens tapes aureus parthenia indistincta → chrysallida indistincta parthenia interstincta → chrysallida obtusa parthenia spiralis → chrysallida spiralis parthenina eximia → chrysallida eximia parvicardium exiguum (gmelin 1791) cardium exiguum parvicardium minimum (philippi 1836) cardium minimum parvicardium ovale (sowerby 1840) cardium fasciatum parvicardium scabrum (philippi 1844) cardium nodosum patella vulgata linnaeus 1758 patina pellucida → helcion pellucidum pecten groenlandicus → palliolum greenlandicum pecten islandicus → chlamys islandica pecten maximus (linnaeus 1758) pecten opercularis → aequipecten opercularis pecten septemradius → pseudamussium septemradiatum pecten similis → similipecten similis pecten varius → chlamys varia pelseneeria stylifera (turton 1825) peringia ulvae → hydrobia ulvae phaxas pellucidus (pennant 1777) cultellus pellucidus solen pellucidus philbertia purpurea → raphitoma purpurea philine aperta (linnaeus 1767) philine aperta quadripartita philine catena (montagu 1803) philine denticulata (adams 1800) philine punctata (adams 1800) philine quadrata (s. wood 1839) philine scabra (müller 1776) philine aperta quadripartita → philine aperta philinoglossa helgolandica hertling 1932 pholas candida → barnea candida pholas dactylus linnaeus 1758 plagiocardium papillosum poli 1795 cardium papillosum pododesmus patelliformis (linnaeus 1761) anomia patelliformis pododesmus squama (gmelin 1791) polygireulima monterosatoi (monterosato 1890) polygireulima sinuosa (sacco 1836) portlandia arctica (gray 1824) yoldia arctica portlandia frigida → yoldiella frigida portlandia lenticula → yoldiella lenticula portlandia lucida → yoldiella lucida portlandia tenuis → yoldiella philippiana potamopyrgus antipodarum (gray 1853) psammobia faeroeensis → gari fervensis psammobia vespertina → gari depressa pseudamussium septemradiatum (müller 1776) pecten septemradius psiloteredo megotara (forbes & hanley 1848) puncturella noachina (linnaeus 1771) purpura lapillus → nucella lapillus putilla semistriata → cingula semistriata raphitoma asperrima (brown 1827) raphitoma leufroyi (michaud 1821) raphitoma linearis (montagu 1803) claturella linearis raphitoma purpurea (montagu 1803) philbertia purpurea retusa obtusa (montagu 1803) utriculus obtusus utriculus pertenuis retusa truncatula (bruguière 1792) utriculus truncatulus utriculus mammillatus retusa umbilicata (montagu 1803) utriculus umbilicatus utriculus nitidulus rhizorus acuminatus (bruguière 1792) rissoa abyssicola → alvania abyssicola rissoa albella lovén 1846 turboella albella rissoa cimicoides → alvania cimicoides rissoa inconspicua alder 1844 turboella inconspicua rissoa interrupta → rissoa parva rissoa jan mayeni → alvania jan mayeni rissoa lactea → alvania lactea rissoa lilacina → rissoa violacea rissoa membranacea (j. adams 1800) rissoa parva (da costa 1779) turboella interrupta rissoa interrupta rissoa punctura → alvania punctura rissoa scrobiculata → alvania scrobiculata geus bulletin no 3.pmd 28-06-2004, 08:46185 186 rissoa semistriata → cingula semistriata rissoa striata → onoba semicostata rissoa violacea desmarest 1814 rissoa lilacina rissoa vitrea → onoba vitrea saxicava arctica → hiatella arctica saxicava rugosa → hiatella rugosa saxicavella jeffreysi winckworth 1930 turneria jeffreysi arcinella plicata scalaria communis → epitonium clathrus scalaria turtonae → epitonium turtonis scaphander lignarius (linnaeus 1758) scaphander punctostriatus (mighels & adams 1841) scissurella crispata fleming 1828 scrobicularia piperata → scrobicularia plana scrobicularia plana (da costa 1778) scrobicularia piperata serripes groenlandicus (bruguière 1798) cardium groenlandicum similipecten similis (laskey 1811) pecten similis siphonodentalium lobatum (sowerby 1860) skenea basistriata (jeffreys 1877) skenea planorbis → skeneopsis planorbis skenea serpuloides (montagu 1808) skeneopsis planorbis (fabricius 1780) skenea planorbis solecurtus chamasolen (da costa 1778) solecurtus scopula (turlok 1822) solen ensis → ensis ensis solen pellucidus → phaxas pellucidus spiralis balea → limacina retroversa spiralis retroversus → limacina retroversa spisula elliptica (brown 1827) spisula solida (linnaeus 1758) mactra solida spisula subtruncata (da costa 1778) mactra subtruncata stenotis palidula → lacuna pallidula syndesmya alba → abra alba syndesmya ovata → abra segmentum syndesmya prismatica → abra prismatica syrnola laevis → eulimella laevis tapes aureus → paphia aurea senescens tapes aureus → paphia aurea tapes decussatus (linnaeus 1758) tapes edulis → venerupis rhomboides tapes pullastra → venerupis pullastra tapes senescens → paphia aurea senescens tapes virgineus → venerupis rhomboides taranis borealis bouchet & warén 1980 taranis moerchi (malm 1861) tectura virginea → acmaea virginea tellimya ferruginosa (montagu 1803) montacuta ferruginosa tellina balthica → macoma balthica tellina calcaria → macoma calcarea tellina crassula → macoma torelli tellina donacina linnaeus 1758 tellina fabula → fabulina fabula tellina loveni → macoma loveni tellina pusilla → tellina pygmaea tellina pygmaea (lovén 1846) tellina pusilla tellina tenuis → angulus tenuis teredo navalis linnaeus 1758 theodoxus fluviatilis (linnaeus 1758) neritina fluviatilis thracia convexa (wood 1815) thracia gracilis (jeffreys 1865) thracia papyracea → thracia phaseolina thracia phaseolina (lamarck 1818) thracia papyracea thracia villosiuscula (macgillivray 1827) thyasira croulinensis (jeffreys 1847) thyasira flexuosa (montagu 1803) axinus flexuosus thyasira sarsi (philippi 1845) timoclea ovata (pennant 1777) venus ovata tonicella marmorea (fabricius 1780) tonicella rubra (linnaeus 1767) tridonta borealis schumacher 1817 astarte borealis tridonta elliptica (brown 1827) astarte compressa tridonta montagui (dillwyn 1817) astarte banksie triforis perversa → triphora adversa triforis perversa adversa → triphora adversa triphora adversa (montagu 1803) triforis perversa triforis perversa adversa trivia arctica (pulteney 1799) trivia monacha (da costa 1778) cypraea europaea trochus cineraria → gibbula cineraria trochus tumida → gibbula tumida trophon clathratus → boreotrophon clathratus trophonopsis barvicensis (johnston 1825) troschelia bernicensis (king 1846) geus bulletin no 3.pmd 28-06-2004, 08:46186 187 turboella albella → rissoa albella turboella inconspicua → rissoa inconspicua turboella interrupta → rissoa parva turbonilla acuta → turbonilla delicata turbonilla crenata (brown 1827) turbonilla rufa turbonilla delicata (monterosato 1874) turbonilla acuta turbonilla lactea (linné 1758) chemnitzia lactea turbonilla rufa → turbonilla crenata turbonilla sinuosa (jeffreys 1884) turneria jeffreysi → saxicavella jeffreysi turrisipho moebii (dunker & metzger 1874) turritella communis risso 1826 turritella terebra turritella erosa couthouy 1838 turritella terebra → turritella communis turtonia minuta (fabricius 1780) cyamium minutum utriculus mammillatus → retusa truncatula utriculus nitidulus → retusa umbilicata utriculus obtusus → retusa obtusa utriculus pertenuis → retusa obtusa utriculus truncatulus → retusa truncatula utriculus umbilicatus → retusa umbilicata velutina plicatilis (müller 1776) velutina velutina (müller 1776) venerupis pullastra (montagu 1803) tapes pullastra venerupis rhomboides (pennant 1777) tapes edulis tapes virgineus venus fasciata → clausinella fasciata venus gallina → chamelea striatula venus ovata → timoclea ovata vitreolina collensi (sykes 1903) vitreolina philippii (rayneval & ponzi 1854) eulima distorta balcis devians xylophaga dorsalis turton 1822 yoldia arctica → portlandia arctica yoldia hyperborea lovén 1859 yoldiella frigida (torell 1859) portlandia frigida yoldiella lenticula (möller 1842) portlandia lenticula yoldiella lucida (lovén 1846) portlandia lucida yoldiella philippiana (nyst 1845) portlandia tenuis zirfaea crispata (linnaeus 1758) geus bulletin no 3.pmd 28-06-2004, 08:46187 188 index of species a abra alba 15, 17, 19, 20, 22, 23, 87, 105, 113, 114, 116, 121, 123, 134, 135, 139, 142, 145, 148, 150, 160, 161, 162, 168, 180, 186, 201, 204, 209, 213 abra nitida 19, 20, 22, 23, 87, 111, 123, 124, 129, 142, 145, 148, 150, 156, 157, 180, 201, 205, 209, 213 abra prismatica 19, 22, 23, 87, 88, 106, 113, 114, 119, 121, 123, 142, 148, 150, 166, 180, 186, 201, 205, 209, 213 abra segmentum 88, 113, 114, 121, 172, 180, 186, 205, 213 acanthocardia echinata 17, 18, 20, 22, 23, 78, 105, 113, 114, 118, 119, 120, 121, 123, 139, 141, 143, 145, 148, 150, 180, 181, 201, 204, 208, 213 acanthocardia-venerupis 155 acar nodulosa 66, 180, 200, 208 aclis ascaris 19, 40, 144, 180, 199, 202, 207, 212 aclis minor 18, 19, 22, 41, 104, 120, 141, 144, 150, 167, 180, 199, 202, 207, 212 aclis walleri 19, 41, 144, 180, 199, 202, 207, 212 acmaea tessulata 17, 26, 140, 165, 180, 198, 202, 205, 210 acmaea virginea 17, 21, 26, 140, 147, 162, 165, 180, 186, 198, 202, 206, 211 acteon tornatilis 18, 20, 21, 55, 121, 123, 124, 141, 144, 148, 166, 180, 199, 203, 208, 212 adipicola simpsoni 69, 180, 200, 208 admete viridula 50, 127, 180, 199, 203, 205, 210 aequipecten opercularis 18, 21, 69, 121, 141, 148, 180, 185, 200, 204, 208, 212 akera bullata 15, 17, 18, 21, 59, 113, 114, 121, 134, 138, 141, 148, 162, 165, 180, 200, 203, 208, 212 alvania abyssicola 31, 123, 180, 185, 198, 202, 207, 211 alvania cimicoides 21, 31, 147, 149, 167, 180, 185, 202, 211 alvania cruenta 32, 127, 172, 180, 202, 209 alvania jan mayeni 32, 127, 185, 202, 210 alvania jeffreysi 31, 180, 198, 207 alvania lactea 17, 21, 31, 142, 143, 148, 149, 165, 166, 180, 185, 202, 213 alvania punctura 17, 21, 32, 141, 147, 180, 185, 198, 202, 207, 211 alvania scrobiculata 32, 127, 180, 185, 202, 210 amauropsis islandicus 38, 180, 198, 205 amphilepis norvegica 126, 156 angulus fabulus 182 angulus tenuis 15, 19, 20, 22, 23, 83, 111, 113, 121, 134, 135, 142, 145, 148, 150, 180, 186, 201, 204, 209, 213 anomia aculeata 182 anomia ephippium 18, 22, 72, 142, 148, 180, 204, 213 anomia patelliformis 180, 185 anomia squamula 180, 182 antalis agile 61, 180, 200, 208 antalis entalis 61, 125, 180, 182, 200, 203, 206, 211 antalis eutalis 126 aporrhais pespelicani 16, 18, 19, 21, 22, 36, 108, 113, 116, 120, 123, 135, 136, 141, 144, 147, 150, 160, 162, 165, 180, 198, 202, 207, 212 aporrhais serresianus 36, 180, 198, 207 arca glacialis 180, 181 arca-astarte crenata 133 arcinella plicata 180, 186 arcopagia crassa 83, 180, 201, 209 arctica islandica 10, 15, 17, 19, 20, 22, 23, 88, 101, 111, 112, 114, 116, 120, 122, 130, 131, 132, 134, 135, 138, 140, 143, 144, 147, 149, 154, 172, 180, 182, 201, 205, 206, 211 assiminea grayana 35, 181, 198, 207 astarte banksie 181, 186 astarte borealis 131, 181, 186 astarte compressa 181, 186 astarte sulcata 77, 181, 201, 207 axinopsida orbiculata 74, 127, 128, 130, 181, 204, 210 axinopsis orbiculata 181 axinus ferruginosus 181 axinus flexuosus 181, 186 b balanus balanoides 131 balanus crenata 128 balcis devians 181, 187 barleeia unifasciata 22, 31, 150, 181, 202, 211 geus bulletin no 3.pmd 28-06-2004, 08:46188 189 barnea candida 15, 19, 20, 22, 23, 94, 106, 113, 121, 134, 142, 145, 148, 150, 181, 185, 201, 205, 209, 213 bathyarca glacialis 66, 101, 127, 128, 132, 133, 159, 172, 180, 181, 204, 209 bathyarca pectunculoides 66, 181, 200, 206 bela exarata 48, 122, 181, 199, 203, 205, 210 bela incisula 181 bela nobilis 181 bela trevelliana 181 bela turricola 181 bela violacea 181 bittium reticulatum 16, 17, 18, 19, 21, 22, 35, 107, 113, 118, 120, 123, 124, 129, 134, 135, 136, 138, 141, 144, 147, 150, 157, 160, 162, 164, 171, 181, 198, 202, 207, 212 bittium-varicorbula 155 boreotrophon clathratus 44, 130, 131, 155, 181, 186, 199, 202, 205, 210 boreotrophon truncatus 44, 181, 199, 206 brachystomia carozzai 50, 181, 199, 208 brachystomia eulimoides 17, 18, 19, 21, 50, 121, 138, 141, 144, 147, 181, 184, 199, 203, 208, 212 brissopsis lyrifera 76 buccinum cyaneum 45, 130, 181, 202, 210 buccinum groenlandicum 158, 181 buccinum undatum 17, 18, 19, 21, 22, 45, 108, 118, 119, 120, 123, 131, 138, 140, 144, 147, 150, 162, 172, 181, 199, 202, 206, 211 c cadulus jeffreysi 100, 181 cadulus subfusiforme 60, 125, 126, 181, 200, 203, 208, 212 caecum glabrum 18, 19, 21, 35, 113, 120, 141, 144, 147, 162, 165, 181, 198, 202, 207, 212 calliostoma formosa 27, 181, 198, 207 calliostoma zizyphinum 27, 181, 198, 207 callochiton septemvalvis 25, 181, 198, 207 capulus ungaricus 37, 181, 198, 207 cardium edule 181 cardium groenlandicum 181, 186 cardium minimum 181 cardium nodosum 181 cardium norvegicum 181 cardium papillosum 181, 185 cerastoderma edule 15, 16, 17, 18, 20, 22, 80, 113, 114, 116, 118, 121, 134, 136, 139, 141, 143, 145, 148, 154, 159, 160, 161, 162, 181, 201, 204, 209, 213 cerastoderma glaucum 16, 17, 18, 80, 136, 139, 141, 181, 201, 204, 209, 213 cerithiella metula 39, 181, 199, 207 cerithiopsis barleei 18, 39, 141, 181, 199, 202, 207, 212 cerithiopsis tubercularis 18, 40, 121, 141, 181, 202, 212 chamelea striatula 19, 20, 22, 23, 89, 105, 111, 113, 114, 121, 124, 142, 145, 148, 150, 161, 162, 168, 170, 181, 187, 201, 205, 209, 213 chemnitzia lactea 181, 187 chlamys islandica 69, 128, 130, 131, 158, 181, 185, 204, 211 chlamys varia 18, 20, 22, 23, 70, 106, 121, 141, 145, 148, 150, 162, 181, 185, 200, 204, 208, 212 chrysallida eximia 18, 51, 123, 124, 140, 181, 185, 203, 211 chrysallida decussata 18, 23, 51, 109, 141, 150, 165, 181, 199, 203, 208, 212 chrysallida indistincta 18, 19, 21, 51, 141, 144, 147, 165, 181, 185, 199, 203, 208, 212 chrysallida obtusa 18, 21, 51, 113, 121, 141, 147, 181, 185, 199, 203, 208, 212 chrysallida spiralis 17, 18, 19, 21, 52, 113, 114, 118, 119, 121, 138, 141, 144, 147, 181, 185, 199, 203, 208, 212 cingula semistriata 17, 21, 32, 141, 147, 181, 185, 186, 198, 202, 207, 212 cingula turgida 19, 32, 144, 181, 198, 202, 207, 211 circe minima 182 claturella linearis 185 clausinella fasciata 19, 20, 22, 89, 142, 145, 148, 149, 181, 187, 201, 205, 209, 213 clinocardium ciliatum 80, 122, 124, 127, 154, 156, 181, 204, 210 clione limacina 60, 132, 181, 200, 203, 206, 210 cochlodesma praetenue 20, 23, 96, 106, 125, 145, 150, 170, 181, 201, 205, 209, 213 colus gracilis 45, 181, 199, 208 colus jeffreysianus 45, 181, 199, 208 colus sabini 46, 181, 199, 207 corbula gibba 15, 16, 17, 19, 20, 22, 23, 92, 105, 106, 111, 113, 116, 119, 120, 121, 124, 134, 136, 139, 142, 143, 145, 148, 150, 159, 160, 161, 162, 164, 170, 181, 201, 205, 209, 213 crenella decussata 69, 127, 181, 200, 204, 205, 210 crepidula fornicata 36, 37, 181, 198, 207 cultellus pellucidus 181, 185 geus bulletin no 3.pmd 28-06-2004, 08:46189 190 cuspidaria obesa (lovén 1846) 97 cuspidaria cuspidata 97, 181, 201, 209 cyamium minutum 182, 187 cylichna alba 18, 20, 23, 56, 110, 128, 130, 131, 140, 143, 149, 182, 203, 210 cylichna cylindracea 18, 21, 55, 141, 148, 166, 182, 199, 203, 208, 212 cylichna occulta 56, 126, 127, 130, 131, 158, 172, 182, 203, 209 cylichna propinqua 182 cylichna scalpta 182 cypraea europaea 182, 186 cyprina islandica 116, 152, 180, 182 cytharella coarctata 21, 47, 121, 147, 182, 183, 199, 202, 208, 212 d delectopecten vitreus 18, 70, 100, 125, 126, 140, 156, 165, 182, 200, 204, 206, 211 dentalium entalis 180, 182 dentalium vulgare 62, 99, 125, 172, 182, 203, 213 devonia perrieri 77, 182, 201, 208 diaphana hyalina 182 diaphana minuta 18, 21, 58, 140, 146, 182, 200, 203, 206, 210 divaricella divaricata 182 donax vittatus 19, 20, 21, 22, 23, 85, 86, 106, 110, 111, 121, 122, 142, 143, 145, 148, 150, 155, 167, 169, 170, 182, 201, 204, 209, 213 dosinia exoleta 20, 22, 91, 145, 148, 149, 166, 167, 170, 182, 201, 205, 209, 213 dosinia lincta 20, 22, 91, 113, 114, 121, 145, 148, 149, 166, 167, 169, 170, 182, 201, 205, 209, 213 dosinia lupinus 182 e ebala nitidissima 18, 21, 52, 113, 114, 121, 141, 148, 165, 182, 199, 203, 208, 212 echinocardium cordatum 105, 110 emarginula fissura 26, 182, 198, 207 ensis americanus 82 ensis arcuatus 82, 182, 201, 209 ensis ensis 19, 20, 22, 82, 113, 121, 123, 141, 145, 148, 162, 165, 182, 186, 201, 204, 209, 213 ensis siliqua 83, 182, 201, 209 entalina tetragona 61, 102, 125, 126, 182, 203, 212 enteroxenos oestergreni 44, 182, 199, 207 epitonium clathratulum 40, 182, 199 epitonium clathrus 18, 19, 21, 40, 113, 114, 121, 141, 144, 147, 182, 186, 199, 202, 207, 212 epitonium trevelyanum 22, 40, 150, 182, 199, 202, 207, 212 epitonium turtonis 18, 21, 40, 141, 147, 165, 166, 182, 186, 199, 202, 207, 212 eulima bilineata 41, 182, 199, 207 eulima distorta 182, 187 eulimella acicula 182 eulimella laevis 18, 19, 21, 52, 141, 144, 148, 165, 182, 186, 199, 203, 208, 212 eulimella nitidissima 182 eulimella scillae 18, 23, 52, 104, 105, 123, 124, 141, 150, 167, 182, 199, 203, 208, 212 evalea divisa 182 f fabulina fabula 19, 20, 22, 23, 84, 111, 121, 142, 145, 148, 150, 162, 170, 180, 182, 186, 201, 204, 209, 213 fucus serratus 107 fusus antiquus 182 g gari depressa 22, 86, 148, 149, 167, 182, 185, 204, 213 gari fervensis 19, 22, 23, 86, 111, 140, 147, 149, 150, 166, 182, 185, 201, 204, 207, 211 gari tellinella 86, 182, 201, 209 gastrana fragilis 84, 113, 114, 119, 121, 172, 182, 204, 213 gibbula cineraria 17, 19, 21, 27, 120, 141, 144, 147, 157, 182, 186, 198, 202, 207, 211 gibbula tumida 17, 21, 27, 141, 147, 165, 182, 186, 198, 202, 207, 211 glossus humanus 10, 89, 182, 201, 209 gouldia minima 91, 113, 114, 121, 172, 181, 182, 205, 213 graphis albida 22, 42, 43, 109, 150, 182, 202, 212 h haliella stenostoma 41, 182, 199, 206 haminea navicula 182 haminoea navicula 55, 113, 114, 119, 121, 172, 182, 203, 213 hanleya hanleyi 25, 182, 198, 206 helcion pellucidum 17, 21, 26, 141, 147, 165, 166, 182, 184, 185, 198, 202, 207, 211 geus bulletin no 3.pmd 28-06-2004, 08:46190 191 hemiaclis ventrosa 22, 43, 44, 109, 150, 182, 202, 212 heteranomia squamula 17, 18, 20, 22, 23, 72, 106, 113, 120, 138, 140, 142, 144, 147, 150, 154, 180, 182, 200, 204, 206, 211 hiatella arctica 15, 19, 20, 22, 23, 93, 100, 118, 119, 120, 123, 125, 126, 127, 128, 130, 140, 144, 147, 149, 158, 162, 182, 186, 201, 205, 206, 210 hiatella rugosa 17, 22, 93, 138, 147, 182, 186, 201, 205, 206, 210 hinia incrassata 18, 21, 46, 123, 141, 147, 149, 155, 165, 166, 182, 184, 199, 202, 208, 212 hinia pygmaea 18, 19, 21, 23, 46, 47, 106, 108, 113, 118, 120, 121, 141, 144, 147, 150, 182, 184, 199, 202, 208, 212 hinia reticulata 15, 16, 17, 18, 19, 21, 23, 47, 106, 113, 116, 118, 119, 121, 123, 124, 134, 135, 136, 138, 141, 144, 147, 150, 162, 182, 184, 199, 202, 208, 212 hydrobia neglecta 30, 182, 198, 207 hydrobia ulvae 15, 16, 17, 19, 21, 22, 30, 107, 113, 116, 118, 120, 134, 135, 138, 139, 141, 144, 147, 150, 162, 164, 182, 185, 198, 202, 207, 211 hydrobia ventrosa 15, 16, 17, 19, 30, 31, 134, 135, 138, 141, 144, 198, 202, 207, 211 i iothia fulva 17, 26, 141, 182, 198, 202, 207, 211 ischnochiton albus 25, 182, 198 j jujubinus clelandi 27, 182, 198, 207 k kellia suborbicularis 22, 77, 148, 149, 182, 201, 204, 208, 213 kelliella miliaris 88, 100, 124, 125, 182, 201, 205, 209, 213 l lacuna crassior 29, 183, 198, 205 lacuna pallidula 15, 17, 19, 21, 22, 29, 133, 140, 143, 146, 149, 165, 171, 183, 186, 198, 202, 206, 210 lacuna parva 17, 19, 21, 29, 120, 141, 144, 147, 165, 183, 198, 202, 207, 211 lacuna vincta 16, 17, 19, 21, 30, 120, 123, 128, 131, 134, 135, 136, 138, 140, 144, 147, 162, 183, 198, 202, 206, 211 laevicardium crassum 22, 80, 119, 148, 149, 181, 183, 201, 204, 209, 213 lamellaria perspicua 37, 183, 198, 207 leda minuta 184 leda pernula 133, 184 lepeta caeca 26, 183, 198, 205 lepidochitona cinereus 25, 183, 198, 207 leptaxinus ferruginosus 75, 123, 181, 183, 204, 210 leptochiton asellus 25, 183, 198, 205 lepton nitidum 18, 20, 22, 77, 113, 114, 121, 141, 145, 148, 183, 201, 204, 208, 213 lepton squamosum 77, 183, 201, 208 limacina retroversa 59, 100, 123, 125, 126, 130, 131, 183, 186, 200, 203, 206, 210 limaria hians 73, 183, 200, 208 limaria loscombi 73, 200, 208 limatula subauriculata 73, 183, 200, 206 liomesus ovum 46, 183, 199, 207 liostomia clavula 53, 183, 199, 208 littorina littorea 15, 16, 17, 19, 21, 28, 113, 114, 116, 118, 120, 134, 135, 136, 138, 141, 144, 147, 154, 159, 160, 161, 162, 163, 183, 198, 202, 207, 211 littorina mariae 28, 183, 198, 207 littorina obtusata 15, 16, 17, 19, 21, 28, 29, 120, 134, 138, 139, 140, 144, 147, 162, 183, 198, 202, 206, 211 littorina saxatilis 15, 16, 17, 19, 21, 29, 120, 130, 131, 132, 133, 135, 138, 139, 140, 143, 146, 160, 183, 198, 202, 206, 210 littorina tenebrosa 15, 16, 17, 21, 29, 114, 134, 135, 138, 139, 141, 147, 162, 164, 165, 183, 198, 202, 207, 211 lucinella divaricata 73, 113, 114, 119, 121, 172, 182, 183, 204, 213 lucinoma borealis 18, 22, 73, 141, 143, 148, 149, 155, 166, 183, 200, 204, 208, 213 lunatia alderi 18, 19, 21, 22, 38, 104, 105, 108, 116, 118, 119, 120, 123, 124, 125, 141, 144, 147, 150, 162, 183, 184, 198, 202, 207, 212 lunatia catena 18, 21, 38, 141, 144, 147, 162, 183, 184, 198, 202, 207, 212 lunatia montagui 22, 38, 108, 150, 183, 198, 202, 207, 212 lunatia pallida 39, 127, 128, 130, 131, 183, 184, 198, 202, 205, 210 geus bulletin no 3.pmd 28-06-2004, 08:46191 192 lutraria lutraria 18, 22, 81, 141, 148, 149, 165, 166, 170, 183, 201, 204, 209, 213 lymnaea peregra 15, 16, 60, 134, 135, 163, 183, 200, 203, 208, 212 lyonsia arenosa 96, 127, 130, 172, 183, 205, 210 lyonsia norvegica 23, 96, 150, 183, 201, 205, 209, 213 m macoma balthica 15, 16, 17, 19, 20, 22, 85, 113, 114, 116, 118, 120, 131, 132, 134, 135, 137, 138, 140, 144, 147, 159, 160, 161, 162, 164, 168, 169, 170, 171, 172, 183, 186, 201, 204, 207, 211 macoma calcarea 11, 22, 56, 85, 99, 101, 102, 116, 122, 126, 127, 128, 129, 130, 131, 132, 146, 154, 155, 156, 157, 158, 168, 183, 186, 201, 204, 205, 210 macoma loveni 85, 130, 172, 183, 186, 204, 210 macoma torelli 85, 130, 131, 158, 172, 183, 186, 204, 209 mactra solida 183, 186 mactra stultorum 18, 20, 22, 23, 81, 110, 113, 114, 122, 141, 145, 148, 150, 162, 165, 183, 201, 204, 209, 213 mactra subtruncata 183, 186 malletia obtusa 66, 183, 200, 208 mangelia attenuata 48, 183, 199, 208 mangelia brachystoma 23, 49, 106, 123, 124, 150, 183, 199, 203, 208, 212 mangelia costata 182, 183 mangelia nebula 49, 183, 199, 208 margarites helicinus 17, 27, 140, 164, 183, 198, 202, 205, 210 melanella alba 22, 43, 150, 167, 183, 202, 212 melanella lubrica 22, 43, 104, 150, 183, 199, 202, 207, 212 melaraphe neritoides 28, 183, 198, 207 menestho divisa 183 modiola modiolus 183 modiolaria discors 183, 184 modiolaria laevigata 183 modiolaria marmorata 183 modiolaria nigra 183 modiolaria tumida 18, 21, 69, 121, 141, 148, 183, 200, 204, 208, 212 modiolula phaseolina 15, 18, 21, 67, 113, 114, 121, 134, 135, 141, 148, 183, 184, 200, 204, 208, 212 modiolus adriaticus 18, 21, 67, 141, 148, 166, 183, 184, 200, 204, 208, 212 modiolus modiolus 15, 18, 21, 68, 116, 118, 120, 134, 135, 140, 143, 147, 162, 168, 183, 184, 200, 204, 206, 211 montacuta ferruginosa 114, 186 montacuta substriata 75, 184, 201, 208 musculus discors 15, 17, 18, 20, 21, 23, 68, 110, 133, 138, 140, 144, 146, 149, 183, 184, 200, 204, 206, 210 musculus laevigatus 68, 127, 128, 130, 183, 184, 204, 210 musculus niger 68, 122, 127, 128, 130, 131, 183, 184, 200, 204, 205, 210 musculus tumidus 183 mya arenaria 15, 17, 22, 23, 37, 92, 134, 135, 137, 138, 147, 150, 163, 168, 169, 170, 179, 184, 205, 207, 211 mya truncata 15, 17, 19, 20, 22, 92, 112, 114, 116, 120, 123, 126, 127, 128, 130, 133, 135, 138, 140, 144, 147, 158, 159, 161, 184, 201, 205, 206, 210 myrtea spinifera 74, 184, 201, 208 mysella bidentata 15, 17, 18, 20, 22, 23, 75, 103, 105, 110, 113, 116, 120, 123, 134, 135, 138, 140, 144, 147, 150, 162, 167, 183, 184, 201, 204, 206, 211 mysella dawsoni 76, 184, 201, 206 mysella tumidula 75, 184, 201, 209 mysia undata 19, 22, 91, 121, 142, 148, 183, 184, 201, 205, 209, 213 mytilaster lineatus 67, 119, 121, 122, 172, 184, 204, 213 mytilaster solidus 67, 113, 114, 172, 184, 204, 213 mytilus adriaticus 183 mytilus edulis 15, 16, 17, 18, 20, 21, 23, 51, 67, 68, 106, 110, 112, 114, 116, 118, 119, 120, 123, 124, 129, 131, 132, 134, 135, 137, 138, 139, 140, 143, 144, 147, 150, 156, 157, 158, 159, 160, 161, 162, 172, 184, 200, 204, 206, 211 mytilus phaseolinus 114, 183 mytilus-cerastoderma 155 n nacella pellucidum 182 natica affinis 39, 75, 128, 130, 131, 184, 202, 210 neptunea antiqua 21, 46, 147, 149, 182, 184, 199, 202, 208, 212 neptunea despecta 46, 130, 131, 184, 202, 210 geus bulletin no 3.pmd 28-06-2004, 08:46192 193 neritina fluviatilis 184, 186 nototeredo norvegica 95, 184, 201, 209 nucella lapillus 18, 19, 21, 44, 140, 144, 147, 184, 185, 199, 202, 206, 211 nucula nitida 184 nucula nitidosa 18, 20, 21, 23, 62, 110, 113, 114, 116, 118, 121, 141, 145, 148, 150, 162, 184, 200, 203, 208, 212 nucula nucleus 18, 20, 21, 23, 62, 110, 116, 123, 141, 145, 148, 150, 165, 184, 200, 203, 208, 212 nucula sulcata 20, 63, 113, 114, 121, 123, 124, 145, 184, 200, 204, 208, 212 nucula tenuis 184 nuculana minuta 23, 63, 102, 105, 120, 123, 124, 132, 133, 149, 172, 183, 184, 200, 204, 206, 211 nuculana pernula 64, 99, 100, 102, 122, 125, 126, 127, 129, 130, 131, 132, 155, 183, 184, 200, 204, 205, 210 nuculoma hanleyi 63, 184, 200, 208 nuculoma tenuis 18, 20, 63, 123, 126, 127, 128, 130, 131, 140, 143, 158, 184, 200, 204, 206, 210 o obtusella alderi 32, 184, 198, 207 ocenebra erinacea 44, 184, 199, 209 odostomia acuta 18, 53, 141, 184, 199, 203, 208, 212 odostomia albella 18, 20, 21, 54, 121, 141, 144, 148, 184, 199, 203, 208, 212 odostomia conoidea 15, 18, 19, 21, 23, 53, 134, 135, 141, 144, 148, 150, 184, 199, 203, 208, 212 odostomia eulimoides 184 odostomia insculpta 184 odostomia pallida 184 odostomia plicata 18, 21, 54, 141, 148, 184, 199, 203, 208, 212 odostomia rissoides 184 odostomia scalaris 18, 21, 51, 113, 121, 141, 147, 184, 199, 203, 208, 212 odostomia turrita 18, 21, 54, 123, 141, 148, 184, 199, 203, 208, 212 odostomia umbilicaris 23, 54, 104, 150, 184, 203, 212 odostomia unidentata 184 oenopota incisula 47, 122, 125, 127, 181, 184, 203, 210 oenopota trevelliana 48, 123, 181, 184, 199, 203, 206, 210 oenopota turricola 18, 21, 23, 48, 130, 131, 140, 146, 149, 181, 184, 199, 203, 205, 210 oenopota violacea 48, 122, 181, 184, 203, 210 omalogyra atomus 15, 17, 18, 21, 50, 134, 135, 138, 140, 147, 182, 184, 199, 203, 206, 211 ondina diaphana 18, 19, 53, 141, 144, 165, 184, 199, 203, 208, 212 ondina divisa 18, 52, 123, 141, 165, 182, 183, 184, 199, 203, 208, 212 onoba aculeus 33, 184, 198, 206 onoba proxima 17, 33, 142, 143, 165, 181, 184, 202, 213 onoba semicostata 16, 17, 21, 33, 134, 135, 138, 139, 140, 147, 162, 164, 181, 184, 186, 198, 202, 206, 211 onoba vitrea 16, 17, 19, 21, 22, 33, 104, 105, 120, 138, 141, 144, 147, 150, 167, 181, 184, 186, 198, 202, 207, 212 ostrea edulis 15, 17, 18, 20, 22, 23, 72, 110, 113, 118, 119, 121, 135, 139, 141, 143, 145, 148, 150, 160, 161, 162, 163, 164, 168, 169, 177, 184, 200, 204, 208, 212 ovatella myosotis 60, 184, 200, 208 p palliolum greenlandicum 70, 127, 128, 129, 184, 185, 204, 210 palliolum striatum 18, 71, 141, 165, 184, 200, 204, 208, 212 palliolum tigerinum 18, 71, 141, 165, 185, 200, 204, 208, 212 pandora glacialis 95, 127, 130, 172, 182, 185, 205, 210 panomya arctica 94, 128, 185, 201, 205, 206, 211 paphia aurea 15, 17, 19, 20, 22, 89, 113, 114, 119, 134, 135, 139, 142, 143, 145, 148, 152, 160, 161, 162, 163, 168, 169, 185, 186, 205, 213 paphia senescens 89, 113, 119, 121, 172, 205, 213 parthenia indistincta 181 parthenia interstincta 114, 181 parvicardium exiguum 15, 16, 17, 18, 20, 22, 78, 113, 118, 121, 134, 136, 139, 141, 145, 148, 160, 162, 181, 185, 201, 204, 209, 213 parvicardium minimum 20, 23, 79, 103, 104, 106, 123, 145, 150, 181, 185, 201, 204, 209, 213 parvicardium ovale 15, 18, 20, 22, 79, 120, 123, 134, 135, 140, 144, 147, 169, 181, 185, 201, 204, 207, 211 geus bulletin no 3.pmd 28-06-2004, 08:46193 194 parvicardium scabrum 15, 17, 18, 20, 22, 79, 121, 134, 135, 139, 141, 145, 148, 162, 181, 185, 201, 204, 209, 213 patella vulgata 17, 26, 140, 185, 198, 202, 207, 211 patina pellucida 182 pecten islandicus 181 pecten maximus 22, 71, 148, 149, 167, 185, 200, 204, 208, 212 pecten septemradius 185 pecten similis 185, 186 pecten varius 181, 185 pecten vitreus 126, 156 pelseneeria stylifera 44, 185, 199, 207 peringia ulvae 185 phaxas pellucidus 19, 20, 22, 23, 83, 105, 110, 114, 121, 123, 141, 145, 148, 150, 159, 170, 181, 185, 186, 201, 204, 209, 213 philbertia purpurea 185 philine aperta 18, 21, 56, 113, 121, 141, 148, 185, 199, 203, 208, 212 philine catena 57, 125, 185, 199, 203, 208, 212 philine denticulata 57, 185, 199, 208 philine punctata 18, 21, 57, 141, 148, 185, 199, 203, 208, 212 philine quadrata 57, 185, 199, 206 philine scabra 57, 185, 199, 208 philinoglossa helgolandica 58, 185, 199, 208 pholas candida 181, 185 pholas dactylus 19, 22, 23, 94, 95, 111, 142, 148, 150, 165, 166, 185, 201, 205, 209, 213 plagiocardium papillosum 79, 121, 122, 155, 172, 181, 185, 204, 213 pododesmus patelliformis 18, 22, 71, 141, 148, 180, 185, 200, 204, 208, 212 pododesmus squama 71, 185, 200, 208 polygireulima monterosatoi 42, 185, 199, 207 polygireulima sinuosa 22, 41, 106, 150, 185, 199, 202, 207, 212 portlandia arctica 11, 50, 64, 100, 101, 120, 126, 127, 128, 129, 130, 131, 132, 133, 154, 156, 157, 158, 159, 172, 185, 187, 204, 209 portlandia frigida 65, 66, 133, 185, 187 portlandia lenticula 133, 185, 187 portlandia lucida 185, 187 portlandia tenuis 185, 187 potamopyrgus antipodarum 30, 185, 198, 207 psammobia faeroeensis 182, 185 psammobia vespertina 182, 185 pseudamussium septemradiatum 71, 73, 123, 124, 129, 185, 200, 204, 208, 212 psiloteredo megotara 95, 185, 201, 205 puncturella noachina 27, 155, 185, 198, 206 purpura lapillus 184, 185 putilla semistriata 185 r raphitoma asperrima 49, 185, 199, 208 raphitoma leufroyi 49, 185, 199, 208 raphitoma linearis 18, 21, 49, 123, 124, 141, 147, 165, 181, 185, 199, 203, 208, 212 raphitoma purpurea 18, 49, 141, 165, 185, 203, 212 retusa obtusa 15, 17, 18, 20, 21, 58, 127, 130, 133, 138, 140, 143, 146, 164, 165, 185, 187, 200, 203, 205, 210 retusa truncatula 15, 16, 17, 18, 20, 21, 23, 58, 105, 113, 121, 134, 135, 136, 138, 141, 144, 148, 150, 162, 185, 187, 200, 203, 212 retusa umbilicata 18, 20, 21, 23, 59, 113, 114, 123, 124, 141, 144, 148, 150, 185, 187, 200, 203, 208, 212 rhizorus acuminatus 59, 185, 200, 208 rissoa abyssicola 185 rissoa albella 16, 17, 19, 21, 22, 33, 34, 120, 134, 135, 136, 138, 141, 144, 147, 150, 160, 185, 187, 198, 202, 207, 212 rissoa cimicoides 185 rissoa inconspicua 16, 17, 19, 21, 34, 113, 114, 118, 119, 120, 134, 135, 136, 138, 141, 144, 147, 162, 185, 187, 198, 202, 207, 212 rissoa interrupta 185 rissoa jan mayeni 185 rissoa lactea 185 rissoa lilacina 185, 186 rissoa membranacea 16, 17, 19, 21, 34, 113, 114, 120, 134, 135, 136, 138, 141, 144, 147, 162, 163, 185, 198, 202, 207, 212 rissoa parva 17, 21, 34, 113, 114, 120, 123, 124, 141, 147, 162, 185, 187, 198, 202, 207, 212 rissoa punctura 185 rissoa scrobiculata 185 rissoa semistriata 186 rissoa striata 186 rissoa violacea 17, 19, 21, 22, 34, 107, 118, 119, 120, 141, 144, 147, 150, 165, 185, 186, 198, 202, 207, 212 rissoa vitrea 186 geus bulletin no 3.pmd 28-06-2004, 08:46194 195 s saxicava arctica 133, 186 saxicava rugosa 186 saxicavella jeffreysi 19, 20, 22, 23, 93, 111, 121, 142, 145, 148, 150, 180, 186, 187, 201, 205, 209, 213 scalaria communis 186 scalaria turtonae 186 scaphander lignarius 56, 186, 199, 208 scaphander punctostriatus 56, 186, 199, 206 scissurella crispata 25, 28, 118, 120, 186, 202, 211 scrobicularia piperata 186 scrobicularia plana 15, 16, 17, 19, 20, 22, 86, 113, 114, 116, 119, 121, 134, 136, 139, 142, 143, 145, 148, 160, 162, 163, 186, 201, 204, 209, 213 serripes groenlandicus 80, 122, 127, 156, 181, 186, 204, 210 similipecten similis 71, 123, 185, 186, 200, 204, 208, 212 siphonodentalium lobatum 61, 100, 102, 125, 126, 132, 186, 203, 210 skenea basistriata 17, 28, 141, 165, 186, 198, 202, 207, 211 skenea planorbis 186 skenea serpuloides 17, 28, 142, 186, 202, 213 skeneopsis planorbis 16, 17, 21, 31, 138, 140, 147, 186, 202, 211 solecurtus chamasolen 82, 186, 201 solecurtus scopula 82, 186, 201, 209 solen ensis 182, 186 solen pellucidus 185, 186 spiralis balea 186 spiralis retroversus 186 spisula elliptica 18, 20, 56, 81, 113, 114, 140, 144, 147, 186, 201, 204, 207, 211 spisula solida 18, 20, 22, 81, 119, 141, 145, 148, 183, 186, 201, 204, 209, 213 spisula subtruncata 15, 19, 20, 22, 23, 81, 82, 104, 107, 110, 113, 116, 118, 119, 121, 134, 141, 143, 145, 148, 150, 161, 162, 163, 170, 183, 186, 201, 204, 209, 213 stenotis palidula 186 syndesmya alba 186 syndosmya prismatica 114 syrnola laevis 182 t tapes decussatus 15, 17, 19, 20, 22, 90, 121, 134, 135, 139, 142, 143, 145, 148, 160, 161, 163, 164, 168, 186, 205, 213 tapes edulis 186, 187 tapes pullastra 186, 187 tapes senescens 186 tapes virgineus 186, 187 taranis borealis 50, 186, 199, 207 taranis moerchi 50, 186, 199, 208 tectura virginea 186 tellimya ferruginosa 18, 20, 22, 23, 76, 105, 110, 113, 114, 120, 140, 144, 147, 150, 162, 167, 183, 186, 201, 204, 206, 211 tellina balthica 186 tellina calcaria 186 tellina crassula 186 tellina donacina 84, 121, 186, 204, 213 tellina fabula 182, 186 tellina loveni 186 tellina pusilla 186 tellina pygmaea 23, 84, 111, 150, 186, 201, 204, 213 tellina tenuis 180, 186 teredo navalis 95, 186, 201, 209 theodoxus fluviatilis 16, 19, 28, 138, 139, 144, 184, 186, 198, 202, 207, 211 thracia convexa 96, 186, 201, 209 thracia gracilis 97, 186, 201, 209 thracia papyracea 186 thracia phaseolina 19, 20, 22, 23, 96, 97, 113, 121, 142, 145, 148, 150, 162, 186, 201, 205, 209, 213 thracia villosiuscula 97, 121, 186, 201, 205, 209, 213 thyasira croulinensis 74, 186, 201, 206 thyasira flexuosa 18, 20, 22, 23, 74, 120, 130, 131, 140, 144, 147, 149, 181, 186, 201, 204, 206, 210 thyasira sarsi 75, 186, 201, 205 timoclea ovata 19, 20, 22, 23, 90, 111, 113, 114, 119, 121, 142, 145, 148, 150, 162, 186, 187, 201, 205, 209, 213 tonicella marmorea 25, 131, 172, 186, 198, 202, 206, 211 tonicella rubra 25, 186, 198, 207 tridonta borealis 15, 18, 77, 116, 127, 128, 130, 131, 133, 135, 138, 140, 143, 168, 181, 186, 201, 204, 205, 210 geus bulletin no 3.pmd 28-06-2004, 08:46195 196 tridonta elliptica 78, 116, 120, 122, 127, 138, 168, 181, 186, 201, 204, 205, 210 tridonta montagui 78, 116, 120, 128, 181, 186, 201, 204, 206, 210 triforis perversa 186 triforis perversa adversa 186 triphora adversa 15, 17, 18, 19, 21, 39, 113, 121, 134, 138, 141, 144, 147, 162, 186, 199, 202, 212 trivia arctica 37, 186, 198, 207 trivia monacha 21, 37, 148, 149, 167, 182, 186, 202, 213 trochus cineraria 182 trochus tumida 182 trophonopsis barvicensis 45, 186, 199, 208 troschelia bernicensis 47, 186, 199, 208 turboella albella 185, 187 turboella inconspicua 185, 187 turboella interrupta 185, 187 turbonilla acuta 187 turbonilla crenata 18, 20, 54, 113, 114, 121, 141, 144, 187, 199, 203, 208, 212 turbonilla delicata 18, 20, 23, 55, 141, 144, 150, 187, 199, 203, 208, 212 turbonilla lactea 18, 20, 21, 55, 113, 114, 118, 121, 141, 144, 148, 181, 187, 199, 203, 208, 212 turbonilla rufa 114, 187 turbonilla sinuosa 23, 150, 187, 203, 212 turneria jeffreysi 186, 187 turrisipho moebii 46, 187, 199, 207 turritella communis 18, 19, 21, 22, 35, 36, 105, 116, 118, 119, 120, 123, 124, 125, 139, 141, 144, 147, 150, 151, 154, 156, 168, 187, 198, 202, 207, 212 turritella erosa 36, 124, 127, 128, 156, 157, 172, 187, 202, 209 turritella terebra 100, 154, 156, 157, 187 turtonia minuta 18, 22, 23, 76, 106, 140, 147, 150, 182, 187, 204, 211 u utriculus mammillatus 185, 187 utriculus nitidulus 185, 187 utriculus obtusus 185, 187 utriculus pertenuis 185, 187 utriculus truncatulus 185, 187 utriculus umbilicatus 185, 187 v velutina plicatilis 37, 187, 198, 206 velutina velutina 37, 187, 198, 206 venerupis pullastra 15, 17, 19, 20, 22, 90, 91, 121, 134, 135, 139, 142, 143, 145, 148, 160, 161, 162, 163, 169, 186, 187, 201, 205, 209, 213 venerupis rhomboides 19, 22, 90, 142, 148, 149, 161, 166, 186, 187, 201, 205, 209, 213 venus fasciata 187 venus gallina 114, 181, 187 venus ovata 186, 187 vitreolina collensi 22, 42, 151, 167, 171, 187, 202, 213 vitreolina philippii 18, 21, 22, 42, 109, 123, 141, 147, 150, 165, 181, 182, 187, 199, 202, 207, 212 x xylophaga dorsalis 95, 187, 201, 209 y yoldia arctica 185, 187 yoldia hyperborea 56, 64, 100, 101, 127, 129, 132, 133, 187, 204, 210 yoldiella frigida 65, 100, 123, 125, 126, 128, 132, 185, 187, 204, 210 yoldiella lenticula 65, 102, 122, 127, 130, 131, 132, 185, 187, 204, 210 yoldiella lucida 65, 123, 185, 187, 200, 204, 206, 211 yoldiella nana 65, 66 yoldiella philippiana 66, 123, 185, 187, 204, 212 z zirfaea crispata 15, 19, 20, 22, 94, 112, 114, 122, 129, 130, 131, 132, 134, 140, 144, 147, 172, 187, 201, 205, 206, 211 geus bulletin no 3.pmd 28-06-2004, 08:46196 geological survey of denmark and greenland bulletin 23, 2011, 81–84 81 on 2 december 2002, eu commissioner poul nielson on behalf of the european development fund signed a €40 million grant to the ghana government. the purpose of this grant was to finance a mining sector support programme (mssp) that covered a broad spectrum of geoscientific projects and other projects aimed at an overall strengthening and modernisation of ghana’s mining sector. one of the major components was collection and interpretation of airborne geophysical data contracted to the two commercial geophysical companies fugro airborne surveys and geotech airborne ltd. the geological survey of denmark and greenland (geus) was contracted to perform the quality control (qc) of the airborne geophysical data collection and processing in a separate mssp project (no 8 acp gh 027/37). the initial provision of quality-control services to the airborne geophysical survey required geus to be on site in ghana for 22 man-months; an expansion of the programme and various circumstances (see below) resulted in a total of 37.25 man-months before the project was completed in january 2010. the danish national space center was subcontracted by geus to perform part of the qc of gravity data acquisition and processing. the qc project was reported by thorning et al. (2010). results from some of the geoscientific projects of the mssp were presented at a workshop in 2008 and a series of small articles from the presentations were published (kalsbeek 2008). organisational setup for the airborne geophysics projects within the mssp interactions between geus and several organisations were required in order to carry out the project. the geological survey department (gsd) within the ministry of lands and national resources of ghana was the main beneficiary institution of the airborne geophysical surveying. a gsd employee acted as supervisor of the projects to collect airborne geophysical data as well as of the qc project. in reality three persons were appointed during the four-year course of the project. other organisations involved were the mssp programme management unit established under a separate contract and supervised by the minerals commission. this is the main promotional and regulatory body for the minerals sector in ghana, which acted as executing organisation for the mssp. in addition, the european delegation in ghana and the national authorising officer of the ministry of finance of ghana took part in the administration. contributions to the projects came from several of the two geophysical companies’ offices; fugro’s offices in accra, johannesburg, ottawa, perth and london and geotech’s offices in accra, toronto and johannesburg were involved. the eu delegation and the national authorising officer undertook the contracting for the various mssp projects. geus did not have formal obligations with respect to the technical specifications in the contract with fugro airborne surveys, but a considerable amount of assistance from geus was required in order to clarify various technical issues. the contract with fugro was signed before geus was awarded the qc project. geus assisted in setting up the tender documents for the airborne data collection project that was later awarded to geotech airborne ltd. in 2008. one of the lessons learned with respect to organising airborne geophysical survey projects similar to those performed in ghana is that the qc team selected for the external quality control should be involved at an early stage and should be consulted concerning the setup of the tender specifications used for the contract with the geophysical consultant performing the measurements. the geophysical survey data the surveys performed by fugro involved the following: (1) remote sensing interpretation (in co-operation with british geological survey (bgs)), (2) acquisition and interpretation of airborne magnetic data, (3) acquisition and interpretation of airborne gamma-spectrometric data, (4) acquisition and interpretation of airborne gravity data and (5) acquisition and interpretation of airborne time-domain electromagnetic (geotem) data. the surveys performed by geotech airborne ltd. involved (1) acquisition and interpretation of airborne magnetic data and (2) acquisition and interpretation of airborne time-domain electromagnetic (vtem) data. during all flights the surveys employed two methods simultaneously, of which magnetic data acquisition was one. quality control of airborne geophysical data from the eu mining sector support programme, ghana thorkild m. rasmussen, leif thorning, arne v. olesen and frands schjøth © geus, 2011. geological survey of denmark and greenland bulletin 23, 81–84. open access: www.geus.dk/publications/bull 8282 volta river basinvolta river basinvolta river basin a1a1a1 a2a2a2 a4a4a4 a5a5a5 a6a6a6a3a3a3 0°1°w 3°w 10°n 9°n 8°n 7°n 6°n 5°n 10°n 9°n 8°n 7°n 6°n 5°n 200 km a8 konongo akwatia hohoe nkwanta th u k keta basin a7 0°1°w3°w 1°e ghana fig. 1. map of ghana with ternary u-th-k images of the gamma-ray spectrometry data from the volta river and keta basin surveys. polygons show survey block boundaries for eight areas covered by detailed geotem surveys (red colour) and for four areas covered by vtem surveys (blue colour). 83 the combined gamma-spectrometry and magnetic survey by fugro included measurements of the horizontal gradient of the magnetic total field in addition to the magnetic total field recordings. the inclusion of the horizontal gradient data improved the lateral resolution in the description of the magnetic field anomalies. the surveys performed by fugro used different fixed-wing aircraft as survey platforms, whereas a helicopter was used for the surveys performed by geotech. figure 1 outlines the areas that were covered by the airborne surveys. the volta river and keta basins were surveyed by fugro. these areas had not previously been covered by detailed airborne geophysical surveys, whereas most of the surrounding ‘basement’ had been covered before. with the completion of the two new surveys the entire onshore area of ghana is now covered by high-resolution magnetic and gamma-spectrometry data. this puts ghana in a leading position with respect to providing modern airborne geophysical data to the mining industry. the volta river and keta basins gamma-spectrometric and magnetic surveys were flown with a 500 m line separation and a survey altitude of 120 m. gravity data were also obtained for the two basin areas, using a mean survey altitude of 860 m above ground and a flight line separation of 5000 m. reconnaissance geotem data were collected over the entire volta river and keta basins using 20 km flight-line separation. later, the areas numbered a1–a8 in fig. 1 were flown with the geotem system using a flight-line separation of 200 m and 400 m. the survey blocks referred to as nkwanta, hohoe, akwatia and konongo were subsequently flown with the vtem system using 400 m line separation. initially, the surveys planned were mainly directed towards obtaining an understanding of the geology of the two basin areas and only included those performed by fugro. vtem transmitter coil vtem receiver coilsvtem receiver coilsvtem receiver coils vtem transmitter bucking coil mag ne to mete r fig. 2. the first author (right) discusses the methods of the vtem system during a break in the surveying of the akwatia block with the geotech operators. the outer transmitter coil has a diameter of c. 26 m and carries a current of c. 200 a before turn-off of the transient signal. 8484 the vtem survey performed by geotech was added to the mssp at a late stage and focused on areas outside the basins. fugro, bgs and geotech performed geological field work as part of a follow-up of the airborne surveying. other mssp geoscientific projects had activities in the areas covered by the geophysical surveys, but the timing of the projects did not allow full integration of data from the various projects. the qc was expanded to include storing of the geophysical data on a server on the geological survey department’s computer system (schjøth et al. 2010). the quality control process carried out by geus was complex and sometimes very difficult, but mostly performed on good terms with the geophysical contractors. the quality of the final data now available in ghana for the mining sector and the scientific community was often significantly improved by the process. qc and training of geological survey department personnel the quality control performed by geus may be viewed as a data assessment independent of the geophysical contractors’ own data quality control. although independent, the quality control by the qc team builds on a high degree of interaction with the geophysical contractors (fig. 2). even though the qc team, the client and the geophysical contractors basically have the same goal of obtaining high-quality data, different views and interests may often exist in terms of defining a proper balance between data quality and project delays. analyses of data with respect to quality from a technical point of view are clearly a major concern of the qc team. some of these analyses follow fairly standardised methodologies and checking procedures, whereas others require an in-depth understanding of data acquisition and processing techniques. in some cases, the acquired field data may be in accordance with the specifications and pass the first routine check, but subsequent application of more advanced checking procedures after the contractor’s processing of the data may reveal problems that were not initially identified. an important part of the obligations by geus was the inclusion of a training component involving two gsd employees and the production of a qc manual (rasmussen et al. 2010) including examples of data issues dealt with during the project. the airborne geophysical contractors also trained gsd personnel. the contracts with geus and the two geophysical contractors were very ambitious with respect to involvement of gsd personnel in the projects. significant knowledge transfer and interaction related to qc took place throughout the entire project period, through regular courses and especially intensive hands-on training provided by geus as part of the actual work with qc. conclusions for obvious reasons the authors are somewhat subjective in assessing the impact of the work performed. nevertheless, we conclude that the geus contribution did have significant influence on the quality of the data released from the project – a conclusion that is supported by statements of gsd personnel, the geophysical contractors and an independent evaluation committee. furthermore, during the course of the project, a constructive working relationship between gsd and geus personnel was established that would be beneficial to both parties in future cooperation. references kalsbeek, f. (ed.) 2008: the voltaian basin, ghana. workshop and excursion, march 10–17, 2008, 136 pp. copenhagen: geological survey of denmark and greenland. http://www.geus.dk/program-areas/common/voltaian_workshop_report.pdf . rasmussen, t.m., thorning, l. & olesen, a.v. 2010: quality control manual for airborne geophysics. the european development fund. project no. 8 acp gh 027/37. mining sector support programme. geophysical investigation. danmarks og grønlands geologiske undersøgelse rapport 2010/42, 207 pp. schjøth, f., rasmussen, t.m. & thorning. l., 2010: guide to the gsd dap-server depository for airborne geophysical data. danmarks og grønlands geologiske undersøgelse rapport 2010/43, 69 pp. thorning, l., rasmussen, t.m. & schjøth, f. 2010: final report. provision of quality control services to the airborne geophysical survey. the european development fund. project no. 8 acp gh 027/37. mining sector support programme. danmarks og grønlands geologiske undersøgelse rapport 2010/44, 35 pp. authors’ addresses t.m.r., l.t.h. & f.s.c., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: tmr@geus.dk. a.v.o., dtu space, national space institute, technical university of denmark, juliane maries vej 30, dk-2100 copenhagen ø, denmark. geological survey of denmark and greenland bulletin 7, 2004, p 13-16 13 seismic data are mainly used to map out structures in the subsurface, but are also increasingly used to detect differences in porosity and in the fluids that occupy the pore space in sedimentary rocks. hydrocarbons are generally lighter than brine, and the bulk density and sonic velocity (speed of pressure waves or p-wave velocity) of hydrocarbon-bearing sedimentary rocks are therefore reduced compared to non-reservoir rocks. however, sound is transmitted in different wave forms through the rock, and the shear velocity (speed of shear waves or s-wave velocity) is hardly affected by the density of the pore fluid. in order to detect the presence of hydrocarbons from seismic data, it is thus necessary to investigate how porosity and pore fluids affect the acoustic properties of a sedimentary rock. much previous research has focused on describing such effects in sandstone (see mavko et al. 1998), and only in recent years have corresponding studies on the rock physics of chalk appeared (e.g. walls et al. 1998; røgen 2002; fabricius 2003; gommesen 2003; japsen et al. 2004). in the north sea, chalk of the danian ekofisk formation and the maastrichtian tor formation are important reservoir rocks. more information could no doubt be extracted from seismic data if the fundamental physical properties of chalk were better understood. the presence of gas in chalk is known to cause a phase reversal in the seismic signal (megson 1992), but the presence of oil in chalk has only recently been demonstrated to have an effect on surface seismic data (japsen et al. 2004). the need for a better link between chalk reservoir parameters and geophysical observations has, however, strongly increased since the discovery of the halfdan field proved major reserves outside four-way dip closures (jacobsen et al. 1999; vejbæk & kristensen 2000). a link between reservoir porosity and sonic velocity in the south arne chalk acoustic properties of the chalk of the danish south arne field have been investigated at three scales by analysing core data, log readings and surface seismic data (japsen et al. 2004). the south arne field is located in the central north sea and chalk porosities of up to 45% are found in the reservoir at almost 3 km depth (fig. 1). the velocity–porosity trend for pure chalk samples from the south arne field matches a modified upper hashin-shtrikman (muhs) curve fitted to ekofisk field chalk (walls et al. 1998). the curve was smoothly extended to 45% porosity using core data from the high-porosity south arne chalk. based on this curve, the acoustic properties of chalk may be calculated as a function of water saturation. this is done by applying gassmann’s equations, which relate the elastic properties of a rock saturated with one fluid to those of the same rock saturated with a different fluid (cf. mavko et al. 1998). the results indicate a pronounced change in the relationship between pand svelocities for chalk saturated with light oil for porosities above c. 30% (fig. 2). this relationship is described by the poisson ratio, and this can provide information about lithology and fluid content of hydrocarbon-bearing reservoirs. for a rock of a given porosity, high pore fluid density results in high p-wave velocity and high poisson ratio, whereas hydrocarbons with low fluid density result in low p-wave velocity and almost unchanged s-wave velocity and a low poisson ratio. these results from the model of the acoustic properties of the chalk indicate that light oil in the high-porous chalk of the south arne field may be detected through amplitude identification of hydrocarbons in chalk reservoirs from surface seismic data: south arne field, north sea peter japsen, anders bruun, ida l. fabricius and gary mavko geological survey of denmark and greenland bulletin 7, 13–16 (2005) © geus, 2005 2820 2900 2900 3000 3000 ekofisk north sea south arne halfdan denmark rigs-1 rigs-2 sa-1 1 km i-1x 56°03´n 4°15´e fig. 1. the south arne field; top chalk structure with location of selected wells. versus offset (avo) inversion of surface seismic data. this is because the variation in seismic reflection amplitude with change in distance between shotpoint and receiver may indicate differences in rock properties above and below the reflector. changes in avo can be directly related to changes in the poisson ratio and thus to differences in pore fluid properties. the uppermost part of the chalk of the south arne field is rich in clay (the ekofisk formation). core data indicate that clayrich chalk has significantly smaller pand s-velocities and a higher poisson ratio than observed for pure chalk (see japsen et al. 2004, fig. 2). the relatively small velocities for a given porosity are probably an artefact due to a reduction in porosity because clay fills up the pore space without affecting the acoustic properties of the chalk matrix. however, the amount of silicates (typically quartz and clay) cannot always be predicted from the gamma log, because the chalk may be rich in very fine-grained silicates which are not all radioactive. in water-wet chalk the amount of very fine-grained material may be estimated from the water saturation (cf. fabricius et al. 2002). the muhs model of the acoustic properties of the chalk is therefore scaled according to silicate content estimated from the water saturation. comparison of well log data and core data the link between the surface seismic data and the reservoir properties in the south arne field is complicated by difficulties in interpreting the sonic log, because mud filtrate has invaded the reservoir near the well bore where the sonic log is registered (fig. 3). the water saturation can be estimated at intermediate and long distances from the bore hole based on resistivity logs, but not adjacent to the well bore. however, the sonic waves travel close to the well bore and it is thus difficult to perform ‘fluid substitution’ of the data, i.e. to transform the acoustic data from the pore fluid environment where they are measured to the fully hydrocarbon-saturated environment of the reservoir (the virgin zone; cf. gommesen 2003). comparison of p-wave velocity and porosity from log and core data clearly shows that the logging data record conditions close to the well bore, where mud filtrate has almost completely flushed the reservoir (fig. 4). core porosities match log porosities estimated from the density log assuming full invasion of mud filtrate (fig. 4a) and p-wave velocities of brine saturated cores generally correspond to readings of pwave velocity in the borehole (fig. 4b). estimation of invasion effects is usually difficult because of the lack of different types of data, and this study underlines the importance of having access to core data. 14 fig. 3. illustration of the effects of invasion of mud filtrate (blue) into a reservoir saturated with hydrocarbons (red). lower left diagram shows how water saturation (sw) increases towards the bore hole. lower right diagram shows how the p-velocity also increases towards the borehole due to the higher density of the mud filtrate compared to the hydrocarbons. the seismic data are influenced by water saturation in the virgin zone (sw, registered by the deep resistivity log), and the acoustic log by the properties close to the well bore, whereas the shallow resistivity log registers the water saturation at some distance from the well bore (sxo). water saturation and sonic velocity are thus not known at the same di-stance from the bore hole, and therefore the acoustic properties of the virgin zone cannot be easily estimated. a b 0.1 0.2 0.3 0.4 0 10 20 30 40 10 20 30 40 50 10 20 30 40 50 1.0 0.8 0.6 0.4 0.2 0.0 m o d u li (g p a) porosity (%) porosity (%) p o is so n r at io ( -) k g sw fig. 2. acoustic properties of chalk as a function of porosity and water saturation (sw). a: bulk (k) and shear modulus (g); b: poisson ratio. note the pronounced difference in the poisson ratio between brine and oil for porosities above c. 30%. muhs prediction based on fine-scale mixing and fluid properties for the south arne field. modified from japsen et al. (2004). w at er s at u ra ti o n distance from drill hole distance from drill hole acoustic log acoustic log seismic data seismic data vp ? sxo ? sw v el o ci ty comparison of well log data and seismic data two approaches may be followed to estimate the acoustic properties of the virgin zone in order to compare well data with seismic data (both are based on gassmann’s equations): 1. transformation of the sonic data to reservoir conditions based on an estimation of the water saturation near the well bore. this can be done using land’s (1968) equation that gives a smooth estimate of the water saturation some distance from the well bore (corresponding to the more scattered registrations by the shallow resistivity log). 2. estimation of the acoustic properties of the reservoir from the muhs model with porosity and water saturation as input (based on the deep resistivity log). estimation of the poisson ratio versus depth in the reservoir depends very much on the approach taken (fig. 4). in the first approach, moderate invasion is assumed and a featureless variation of the poisson ratio results (fig. 4c, dashed black curve). in the second approach, forced displacement of the hydrocarbons near the well bore is assumed and forward modelling results in a characteristic pattern with pronounced peaks at top ekofisk formation and top tor formation, and low values in the highly porous tor reservoir (full red curve). the latter pattern is in good agreement with the inverted seismic data (fig. 5b, blue avo curve). the avo attributes were calculated from the angle-dependent impedance inversions combined with information on the absolute level of chalk velocity which is not contained in the seismic data (bach et al. 2003). acoustic impedance, shear impedance and the poisson ratio were extracted at the location of the rigs-2 well. the avo results show a good match with the well log data. a low poisson ratio in the tor formation near the rigs2 well is in agreement with the presence of light oil in the highly porous chalk of the south arne field. the first approach based on land’s (1968) equation results in a mismatch between log and seismic estimates of the poisson ratio in the virgin zone (fig. 4c, dashed red curve versus fig. 5b, blue avo curve). land’s equation (and the shallow resistivity log) thus apparently underestimates the mud-invasion close to the well bore in the highly porous parts of the reservoir and probably reflects the conditions at some distance from where the sonic log p-wave propagates. this suggestion is further supported by the coincidence of log estimates of density and p-wave velocity with results from core samples saturated with brine. the content of hydrocarbons thus appears to drop to a very low value close to the well bore where the p-wave velocity reaches its maximum value restricting the propagation of p-waves to a very narrow zone, whereas the propagation of s-waves is less affected by the pore fluid content (fig. 3). the second approach based on the muhs model results in a good match between log and seismic estimates of the poisson ratio in the virgin zone (fig. 4c, full red curve versus fig. 5b, blue avo curve). the best way to estimate the acoustic properties of the virgin zone is therefore to use the extended modified upper hashin-shtrikman velocity–porosity relation for chalk. avo inversion of the seismic data based on such synthetic sonic logs reveals a zone of very low 15 muhs sonic log muhs landís eq. a b c m ea su re d d ep th ( m ) 2780 2800 2850 1 2 3 4 0.1 0.2 0.3 0.40 25 50 75 100 sw, porosity (%) velocity (km/s) poisson ratio (-) e ko fis k fm e ko fis k fm vs vp core vp core vs core vp core vs core porosity to r fm to r fm fig. 4. log data and predictions from the corrected muhs model based on porosity and water saturation for the chalk section in the rigs-2 well. a: porosity (estimated from the density log) and water saturation (sw). b: pand s-wave velocity (vp, vs). data (blue curves) and predictions from the corrected muhs model (red curves). c: poisson ratio in the virgin zone; prediction of the corrected muhs model (full red curve), prediction based on land’s (1968) equation (black dashed curve). in the high-porosity oil zone of the tor formation, the oil is predicted to be almost completely flushed. this is indicated by the closeness of the measured vp log (blue curve) and measured vp for the brine saturated cores (green circles), whereas the predicted vp for the virgin zone is low (red curve based on the muhs model and sw). modified from japsen et al. 2004. poisson ratio that correlates with the oil reservoir in the tor formation. in this way avo inversion provides direct evidence for the presence of oil in highly porous chalk of the south arne field. references bach, t., espersen, t.b., pedersen, j.m., rasmussen, k.b., hinkley, r. & pillet, r.p. 2003: seismic inversion of avo data. in: hansen, p.c., jacobsen, b.h. & mosegaard, k. (eds): methods and applications of inversion, 31–42. berlin: springer verlag. fabricius, i.l. 2003: how burial diagenesis of chalk sediments controls sonic velocity and porosity. american association of petroleum geologist bulletin 87, 1755–1778. fabricius, i.l., mavko, g., mogensen, c. & japsen, p. 2002: elastic moduli of chalk as a reflection of porosity, sorting and irreducible water saturation. society of exploration geophysicists technical programme and expanded abstracts 21, 1903–1906. gommesen, l. 2003: prediction of porosity and fluid in chalk from acoustic measurements, 94 pp. ph.d. thesis. lyngby: technical university of denmark. jacobsen, n.l., engstrøm, f., uldall, a. & petersen, n.w. 1999: delineation of hydrodynamic/geodynamic trapped oil in low permeability chalk. society of petroleum engineers paper 56514, 10 pp. japsen, p., bruun, a., fabricius, i.l., rasmussen, r., vejbæk, o.v., pedersen, j.m., mavko, g., mogensen, c. & høier, c. 2004: influence of porosity and pore fluid on acoustic properties of chalk: avoresponse from oil, south arne field, north sea. petroleum geoscience 10, 319–330. land, c.s. 1968: calculation of imbibition relative permeability for twoand three-phase flow from rock properties. society of petroleum engineers journal 1968, 149–156. mavko, g., mukerji, t. & dvorkin, j. 1998: the rock physics handbook, 329 pp. cambridge: cambridge university press. megson, j.b. 1992: the north sea chalk play; examples from the danish central graben. in: hardman, r.f.p. (ed.): exploration britain: geological insights for the next decade. geological society special publication (london) 67, 247–282. røgen, b. 2002: north sea chalk – textural, petrophysical and acoustic properties, 106 pp. ph.d. thesis. lyngby: technical university of denmark. vejbæk, o.v. & kristensen, l. 2000: downflank hydrocarbon potential identified using seismic inversion and geostatistics: upper maastrichtian reservoir unit, dan field, danish central graben. petroleum geoscience 6, 1–13. walls, j.d., dvorkin, j. & smith, b.a. 1998: modeling seismic velocity in ekofisk chalk. society of exploration geophysicists technical programme and expanded abstracts 17, 1016–1019. b tw o w ay t im e (s ) tw o w ay t im e (s ) poisson ratio (-) poisson ratio (-) acoustic impedance (106 kg/m2s) top ekofisk top tor base chalk well log avo result base chalk top ekofisk top tor acoustic impedance (106 kg/m2s) a 2.7 2.8 0.15 0.30 0.45 0.15 0.45 2.7 2.8 5 ne 7 rigs-2 6 8 9 10 sw 134 fig. 5. two-way time section with avo inversion of seismic data and inserted log response for the rigs-2 well computed from forward modelling of the corrected muhs model (fig. 4). a: acoustic impedance (density × vp). b: poisson ratio. very good agreement is observed between avo and log estimates for the acoustic impedance. note the peaks in the tight zones near top ekofisk formation (top chalk) and top tor formation. there is also good agreement between the log and avo estimates of the poisson ratio, e.g. the peak at top tor formation and the low values within the tor formation. this pattern cannot be resolved by the log if the acoustic properties are estimated from the sonic log because the water saturation near the well bore is unknown. seismic quality is severely reduced south-west of the well location due to an overlying gas cap. modified from japsen et al. (2004). 16 authors’ addresses p.j., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: pj@geus.dk a.b., ødegaard & danneskiold-samsøe a/s, titangade 10, dk-2200 copenhagen n, denmark. i.l.f., technical university of denmark, building 115, anker engelundsvej 1, dk-2800 lyngby, denmark. g.m., stanford university, stanford, california 94305-2215, usa. geological survey of denmark and greenland bulletin 33, 2015, 81-84 81 composition of ilmenite and provenance of zircon in northern brazil christian knudsen, tonny b. thomsen, feiko kalsbeek, jeppe a. kristensen, helenice vital and roger k. mclimans th e mineral ilmenite (fetio3) is an important component of heavy-mineral placer deposits and constitutes the largest volume of valuable mineral in such deposits. th e minerals zircon (zrsio4) and rutile (tio2), which occur in lower concentrations than ilmenite in the deposits, have a greater value per ton – c. 1100 and 900 $/ton respectively – compared to ilmenite that ranges from 100 to 200 $/ton depending on its composition. other minerals such as staurolite, sillimanite, amphibole and garnet are generally also present in placer deposits, but are of minor or no commercial value and, e.g. amphibole needs to be separated from the valuable heavy minerals which adds to the production cost. ilmenite is more valuable as a raw material in titanium dioxide manufacture if the titanium content is enhanced by natural leaching of the iron component. when exploring for potentially economic heavy-mineral placer deposits, both the variation in composition and distribution of ilmenite are of interest. accordingly, it is also important to understand not only the concentration of heavy minerals in the ground but also the abundance and composition of the individual minerals. th e source (provenance), route and mechanism of transport from source to potential reservoir sandstones are of interest when attempting to understand petroleum systems in sedimentary basins. heavy minerals in reservoir sandstones contain a wealth of information about their formation that characterises their source. th us a database with the characteristics of possible sediment sources is a key tool to investigate the distribution, composition and other characteristics of the heavy minerals in a given area. both heavy-mineral exploration and locating off shore petroleum reservoir sand sources are relevant in northern brazil. in june 2011, the geological survey of denmark and greenland in cooperation with universidade federal do rio grande do norte, natal, rn, brazil set up a project to sample and analyse in situ cretaceous sandstones, sands in river beds and sands from the coastal sediments. th e results were entered into a database of heavy-mineral compositions and properties. th e variation in the distribution of the heavy minerals as well as their composition were determined for 34 samples from northern brazil using computer-controlled scanning electron microscopy (ccsem; keulen et al. 2012). u-pb ages of detrital zircons were determined by laser ablation inductively coupled plasma mass spectrometry (for analytical procedures see frei et al. 2006). four samples for the u-pb age determinations came from cretaceous sandstones, two from river beds and two from beaches. alteration of the heavy minerals when subjected to alteration in the sedimentary environment, heavy minerals react diff erently depending on the local physical and chemical conditions. in hot and humid conditions minerals such as olivine, pyroxene, amphibole © 2015 geus. geological survey of denmark and greenland bulletin 33, 81–84. open access: www.geus.dk/publications/bull ilmenite leucoxene rutile zircon staurolite silimanite-kyanite garnet epidote amphibole baía de marajó 45°w 3°s rio piranhas-açu amazon basin 200 km brazil são luís parnaíba acaraú !! ! !!! ! !!! ! 40°w45°w 3°s 200 km 62–65 60–62 58–60 56–58 54–56 52–54 tio2 (%) quaternary neogene–quaternary tertiary–quaternary tertiary cretaceous jurassic triassic carboniferous–permian cambrian–devonian neoproterozoic–palaeozoic precambrian litho-chronology belem são luís fortaleza teresina tianguá são luís-grajaû basin parnaíba acaraú camelá sub-basin rio capim baía de marajó fig. 1. modal composition of the heavy-mineral fraction in sand samples from northern brazil. fig. 2. average composition of ilmenite and altered ilmenite in samples from beaches, from river beds and from outcrop of cretaceous sandstone. for location see fig. 1. 8282 and epidote are unstable and gradually disappear from the heavy-mineral assemblage (morton & hallsworth 1999). th e heavy-mineral assemblages in the fi ve easternmost beach sand samples from the northern brazilian coast (fig. 1) all contain abundant amphibole. th e source of those heavy minerals is the precambrian basement in the hinterland (pink, fig. 2). a similar, diverse, heavy-mineral assemblage is described by da silva & vital (2000) in samples from the rio piranhas-açu, north-eastern brazil. th e low degree of alteration of the heavy minerals is probably due to the climate which is arid in this part of brazil. th e heavy-mineral assemblages of the fi ve eastern samples change westwards with increasing contents of alumina-silicates such as staurolite, sillimanite and kyanite; minerals which must be abundant in the hinterland, and which can fi ngerprint the sediment source for the eastern samples. to the west of parnaíba (fig. 1), the heavy-mineral assemblages are dominated by ilmenite, leucoxene, rutile, staurolite and zircon, which are very stable minerals (morton & hallsworth 1999). th e precipitation, humidity and temperature increase towards the west and towards the amazon basin. th is may account for the mineralogical change indicative of intense alteration where even fairly stable heavy minerals like garnet have disappeared. th e cretaceous sandstone in the interior of northern brazil is altered by intense kaolinisation. mendes & truckenbrodt (2009) describe a mature heavy-mineral assemblage from albian sandstones (itapecuru group) in the são luísgrajaû basin and góes et al. (2007) described similar assemblages from the campanian–maastrichtian (ipixuna formation) in the camelá sub-basin to the west. only robust heavy minerals such as ilmenite, zircon, rutile and staurolite are present in the cretaceous samples inland south of são luís and baia de marajá (fig. 1) whereas less stable heavy minerals are lacking. th e higher degree of alteration found in the heavy-mineral assemblage in the coastal beach samples in the western section of the coast could accordingly also be an eff ect of re-deposition of cretaceous sandstones from the hinterland. ilmenite composition th e titanium content of ilmenite changes when it is subjected to chemical weathering; iron is leached and the relative content of titanium increases (fig. 3a). ultimately the mineral leucoxene, which mainly consists of tio2, is formed (bailey et al. 1956). a gradual increase of tio2 in ilmenite in beach sand is evidenced towards the west (fig. 2). as discussed above that may be an eff ect of increasing humidity, but it may also be caused by an infl ux of reworked sediment from altered cretaceous sandstones in the hinterland. th e tio2 content of ilmenite in the albian sandstones (itapetio 2 (%) 50 60 70 80 90 100 ilmenitetitanomagnetite leucoxene rutileb tio 2 (%) 50 60 70 80 90 100 50 40 30 20 10 0 fe 2 o 3 ( % ) ilmenite leucoxene rutile a a b c 50 µm fig. 3. ccsem analysis of titanium minerals from a cretaceous outcrop along rio capim (sample ggu 538118). a: the distribution of tio2 versus fe2o3 in the ilmenite, leucoxene and rutile show an inverse relationship between these two components, where fe2o3 decreases with increasing tio2 – and with the degree of weathering. b: histogram showing the content of tio2 in the titanium mineral grains. fig. 4. scanning electron microscope (sem) backscatter image of an ilmenite grain leached from the rim. in the light grey central part of the grain (a), the ilmenite is un-leached with preserved white hematite lamellae. surrounding the un-leached core there is a zone (b), where the hematite is leached away forming pores (black) and where the ilmenite shows partial leaching (dark grey patches). in the grey rim (c), the ilmenite is highly leached and almost all iron is removed leading to the formation of leucoxene. 83 curu group; eastern line of samples in fig. 2) is high and the tio2 content is even higher in the campanian–maastrichtian (ipixuna formation) located in the more humid climate to the west. th e heavy-mineral assemblage from a cretaceous outcrop along rio capim is dominated by ilmenite. figure 3b shows that the content of tio2 in the titanium minerals has a bimodal distribution with almost no unaltered ilmenite (50% tio2). th e bimodal distribution may indicate that the weathering is a two stage process. an example of chemical weathering of an ilmenite grain is shown in fig. 4. th e grain has a central core that is largely unaltered with hematite lamellae preserved, an intermediate zone where the hematite is leached away giving rise to porosity, and an outer margin that is completely transformed to leucoxene. zircon provenance th e crystalline basement complexes underlying the investigated sedimentary rocks in the eastern part of the area is the borborema province (de almeida et al. 1981) that is composed of a complex assemblage of palaeoproterozoic granitoid and metasedimentary gneisses, locally with outcrops of archaean rocks and with numerous plutons of neoproterozoic granite. neoproterozoic sedimentary successions are also present. palaeoproterozoic granitoid rocks are the most common, forming some 70% of the basement complex. th e são luís craton in the west is dominated by the trans amazonian orogeny with ages ranging from 2000 to 2200 ma (klein et al. 2005). th e u-pb age distribution patterns for eight samples from north-eastern brazil (fig. 5) show the ages of zircon sand grains. four of these are from cretaceous sandstone, two are riverbed sand, and two are coastal sand. th e age distributions (fig. 5) are complex. th e largest component in all samples is formed by palaeofig. 5. u/pb age distributions of detrital zircons from samples from northern brazil. the two green stars to the east represent sandstone samples from the albian itapecuru group and the two green stars to the west represent sandstone samples from the campanian–maastrichtian ipixuna formation. for location see fig. 1. beach sand river bed sand cretaceous sandstone 45°w 3°s 100 km ggu 538254, n = 76/114 ggu 538111, n = 94/140 ggu 538118, n = 86/119 age (ga) 210 3 50 40 30 20 10 0n u m b e r o f a n a ly se s ggu 538205, n = 103/125 age (ga) 210 3 50 40 30 20 10 0n u m b e r o f a n a ly se s ggu 538219, n = 96/119 age (ga) 210 3 50 40 30 20 10 0n u m b e r o f a n a ly se s ggu 538144, n = 75/117 210 3 50 40 30 20 10 0n u m b e r o f a n a ly se s ggu 538125, n = 62/105 age (ga)210 3 50 40 30 20 10 0n u m b e r o f a n a ly se s 30 20 10 0 n u m b e r o f a n a ly se s age (ga)210 3 40 30 20 10 0 age (ga)210 3 age (ga)210 3 50 30 20 10 0n u m b e r o f a n a ly se s 8484 authors’ addresses c.k., t.b.t. f.k. & j.a.k, geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark; e-mail: ckn@geus.dk h.v., universidade federal do rio grande do norte, natal, rn, brazil. r.k.mcl dupont titanium technologies, wilmington, de, usa. proterozoic zircons, 1800–2300 ma, comprising 50–70% of the zircon populations. archaean zircons (2500–3500 ma) constitute 10–25% of the population and neoproterozoic zircons (500–700 ma) constitute 10–20% of the population. only some 15% of the zircons have ages outside these age groups. th ere are no obvious diff erences in age for the zircon populations in the cretaceous sediments, the riverbed sands and the coastal sands. zircons with ages in the range 2000 to 2200 ma, equivalent to the trans-amazonian orogeny (green columns on fig. 5) are common in all samples, in good accordance with observations made by klein et al. (2005). th e content of zircons in the age range 1800–2000 varies considerably. neoproterozoic ages (blue columns on fig. 5) equivalent to braziliano or pan-african orogeny are less frequent in the analysed sands as compared to what is previously described from the borborema province (nascimento et al. 2007). discussion and conclusion th e heavy-mineral assemblages in the east are less mature than assemblages in the west. th at may refl ect a lower degree of alteration of the heavy minerals caused by a drier climate. in the hot and humid area to the west in the amazon basin, the heavy-mineral assemblages are very mature refl ecting the intense chemical attack and removal of heavy minerals such as pyroxene, amphibole and garnet. indications of more intense chemical alteration of the heavy minerals are also seen from the composition of ilmenite that shows decreasing iron content and accordingly increasing content of tio2 towards the west. cretaceous sandstones in the area are kaolinised and the heavy minerals are also strongly altered, most intensely in the hot and humid area in the amazon basin. th e high degree of alteration found in river and beach sediments in the western area could also be caused by reworking of previously altered cretaceous sandstones. apparently, nearly all zircons in the investigated sediments may originate from the underlying crystalline basement, suggesting mainly local source areas. th e zircon-age spectra are fairly uniform suggesting either that the geology in the source area is rather uniform or that the zircons were homogenised in the sedimentary environment and that the sand in the rivers and on the beach at least partly represent reworked cretaceous sandstones. acknowledgements dupont titanium technologies, wilmington, delaware, usa is thanked for fi nancial support. francisco canindé and miguel borges from universidade federal do rio grande do norte, natal, rn, brazil are thanked for help during the fi eldwork. references bailey, s.w., cameron, e.n., spedden, h.r. & weege, r.j. 1956: the alteration of ilmenite in beach sands. economic geology 51, 263–279. da silva, m.g. & vital, h. 2000: provenance of heavy-minerals in the piranhas-açu river, northeastern brazil. brazilian journal of geology 30, 453–456. de almeida, f.f.m., hasui, y., de brito neves, b.b. & fuck, r.a. 1981: brazilian structural provinces: an introduction. earth science reviews 17, 1–29. frei, d., hollis, j.a., gerdes, a., harlov, d., karlsson, c., vasquez, p., franz, c., johansson, l. & knudsen, c. 2006: advanced in situ geochronological and trace element microanalysis by laser ablation techniques. geological survey of denmark and greenland survey bulletin 10, 25–28. góes, a.m., rossetti, d.f. & mendes a.c. 2007: heavy mineral as a tool to refine the stratigraphy of kaolin deposits in the rio capim area, northern brazil. anais da academia brasileira de ciências 79, 457–471. keulen, n.t., frei, d., riisager, p. & knudsen, c. 2012: analysis of heavy minerals in sediments by computer-controlled scanning electron microscopy (ccsem): principles and applications. mineralogical association of canada short course 42, 167–184. klein, e.l., moura, c.a.v. & pinheiro, b.l.s. 2005: paleoproterozoic crustal evolution of the são luís craton, brazil: evidence from zircon geochronology and sm-nd isotopes. gondwana research 8, 177–186. mendes, a.c. & truckenbrodt, w. 2009: provenance of albian sandstones (itapecuru group), eastern border of são luís-grajaú basin, maranhão, using heavy mineral analysis and mineral chemistry. boletim do museu paraense emílio goeldi, ciências naturais 4, 57–74. morton, a.c. & hallsworth c.r. 1999: processes controlling the composition of heavy mineral assemblages in sandstones. sedimentary geology 124, 3–29. nascimento, m.d.s., góes, a.m. macambira, m.j.b. & brod, j.a. 2007: provenance of albian sandstones in the são luís–grajaú basin (northern brazil) from evidence of pb–pb zircon ages, mineral chemistry of tourmaline and and palaeocurrent data. sedimentary geology 201, 21–42. geological survey of denmark and greenland bulletin 23, 2011, 13–16 13 hydrocarbon-bearing upper jurassic sandstone reservoirs at depths of more than 5000 m may form a future exploration target in the danish central graben (fig. 1). the upper jurassic sandstone play in the danish sector has historically been less successful than in the neighbouring norwegian and british sectors of the north sea. this is mainly due to poor reservoir quality of the sandstones. however, the discovery in 2001 of an oil accumulation at a depth of more than 5000 m in the svane-1 well has triggered renewed interest in the upper jurassic high temperature – high pressure (hthp) sandstone play in danish waters. the jurassic plays comprise sandstone reservoirs deposited in a variety of environments, ranging from fluvial to deep marine. this paper presents a study of a minor area around the svane-1a well in the tail end graben (fig. 1). the objective was to map acoustic impedance variations and hence to identify porosity anomalies associated with jurassic sandstone units. interpretation in a tectonic setting such as the jurassic hthp petroleum system in the danish part of the central graben is hampered by low seismic vertical resolution. however, by combining regional seismic mapping with inversion results and petrophysical log analysis, such obstacles can be tackled by mapping acoustic impedance variations. application of seismic inversion techniques for porosity prediction in sandstone is a standard geophysical tool (dolberg et al. 2000). petrophysical analysis of well-log data from the upper part of the jurassic sandstones encountered in the svane1a well shows a relationship between acoustic impedance (ai) and total porosity (phit), see later. this log-derived ai-phit relationship can be applied to transform acoustic impedance variation into porosity variation, when the acoustic impedance is predicted from seismic inversion of a 2d profile, and can be used to locate porosity anomalies associated with sandstone intervals in the area around the svane-1a well. setting the danish central graben is part of the jurassic north sea rift complex and consists of a system of nnw–ssetrending half-grabens bounded by the coffee soil fault and the mid north sea high (fig. 1; japsen et al. 2003; møller & rasmussen 2003). rifting took place from the middle jurassic and persisted into the early cretaceous. the syn-rift sedimentary fill is dominated by mudstone with subordinate layers of sandstone. in some stratigraphic intervals, the mudstone is rich in organic matter (petersen et al. 2010). mapping porosity anomalies in deep jurassic sandstones – an example from the svane-1a area, danish central graben tanni abramovitz norwegian–danish basin c offee soil fault ringkøbing–fyn h igh 4°e 5°e 56°n c entral g raben tail end g raben norway 25 km uk germany heno plateau mid north sea high uk nl n dk 200 km g national border inverted profile svane-1a well high fault fig. 1. map of the danish central graben showing the location of the svane-1a well and the 2d seismic profile that was inverted for acoustic impedance. 5 km 4.2 svane-1a 04°24´36.22´´e inverted profile 56°07´50.56´´n 4.3 4.4 4.5 4.6 4.7 4.8 4.9 5.0 5.1 5.2 tw o -w ay travel tim e (sec) fig. 2. two-way travel time structure map of the intra-kimmeridgian marker horizon corresponding to the top of the drilled svane-1a sandstones. © geus, 2011. geological survey of denmark and greenland bulletin 23, 13–16. open access: www.geus.dk/publications/bull 1414 the svane-1a area the svane-1a well is located on a 4-way-dip closure structure at an intra-kimmeridgian level in the northern part of the tail end graben (fig. 2). it is one of the deepest wells ever drilled in denmark (total depth 5952 m). the structure map of an intra-kimmeridgian marker horizon corresponding to the top of the svane-1a sandstones shows that the well is located on a nnw–sse-oriented structural high along the basin axis bounded by two depocentres (fig. 2). upper jurassic sandstone with dry gas was encountered at 5311 m. unfortunately, no cores or sidewall cores were collected due to unstable borehole walls. at depths over 5400 m, the sandstone layers in svane-1a are characterised by porosities of 15–24% and low permeabilities. the svane-1a well is situated in a hthp environment with overpressures of 8630 psi at a depth of 5350 m, which may imply that the pore pressure is close to the fracture pressure according to johannessen et al. (2010). seismic inversion for acoustic impedance seismic inversion is the process of transforming seismic reflection data into quantitative rock properties such as acoustic impedance (ai) using reflection seismic data constrained by borehole data in order to describe a possible reservoir. acoustic impedance is the product of the rock density and the compressional p-wave velocity, which are both commonly measured in boreholes as the bulk density and the sonic velocity. a log-derived ai-phit relationship based on the petrophysical well log is used to transform the inversionderived ai into total porosity (phit). seismic inversion for acoustic impedance was carried out using the 2d isis seismic inversion software. the inversion algorithm is a deterministic approach based on a simulated annealing algorithm (maver & rasmussen 1995; rasmussen & maver 1996). both seismic and well-log data were used for the inversion. the input data consist of a 2d seismic profile (fig. 3) and raw log data (sonic and density) as well as the time-depth data from the svane-1a well. distance (km) t w o -w ay t ra ve l ti m e (s ec ) random seismic line svane-1a area 0 5 10 15 20 2.5 3.0 3.5 4.0 4.5 5.0 5.5 6.0 se is m ic a m p lit u d e −3000 −2000 −1000 0 1000 2000 3000 enewsw intra-volgian intra-kimmeridgian near base jurassic fig. 3. the input for the seismic inversion is a 2d seismic profile extracted from the 3d pam_99 survey with three interpreted intra-jurassic marker horizons: intra-volgian, intra-kimmeridgian and near base jurassic. the svane-1a well location is indicated by the white line. note the alternating high and low amplitude layers below the intra-kimmeridgian marker horizon. 3.4 3.5 3.6 3.7 3.8 3.9 4.0 4.1 4.2 4.3 4.4 4.5 4.6 4.7 t w o -w ay t ra ve l ti m e (s ec ) 906 1 937 37 968 73 in-line x-line 6.11 7.15 8.26 9.37 10.49 4.08e+06 1.05e+07 well log inversion result low frequency model impedance [106 kg/m2s] 5300 m depth 5628 m t ar ge t in te rv al fig. 4. quality control of the absolute acoustic impedance inversion result. a: section of the inversion result with the acoustic impedance log inserted at the well location. b: comparison between the acoustic impedance trace estimated at the well location (blue), the low frequency model at the well location (green) and the acoustic impedance well log (red). 0.00 4000 6000 8000 10 000 impedance (ai, g cm × msec ) p h it ( fr ac ti o n ) 12 000 14 000 16 000 0.10 0.20 0.30 0.40 0.50 –3 –1 fig. 5. cross plot of the log-derived acoustic impedance (ai) versus the log-derived total porosity (phit), based on the svane-1a sonic and density log data from 5300–5627.93 m. 15 in order to ensure a well-to-seismic tie, the available timedepth data were used to create a reflectivity series with the same sampling rate (4 msec) as the input seismic data and to convert the log data from a depth to a two-way travel time (twt). the acoustic impedance was calculated by multiplying the calibrated density log and the velocity log derived from the calibrated sonic log. the reflectivity series was computed by differentiating the acoustic impedance series. after the log calibration, the svane-1a wavelet was estimated by deriving the convolution operator between the reflectivity log and the seismic trace at the well location using a least squares wavelet estimation method. the length of the wavelet was estimated over the jurassic target interval to 4.2–4.58 sec twt. in general, seismic data have limited frequency bandwidth at the low and high ends. the (missing) low frequencies contain the critical information concerning the absolute values of impedance. in order to invert for absolute acoustic impedance, a low frequency model is needed to introduce the sub-seismic frequencies into the seismic inversion result. the 2d low frequency model is constructed by laterally extrapolating the final calibrated impedance log from the svane-1a well between three interpreted horizons extracted from the 3d seismic pam_99 survey to guide and yield the absolute level of acoustic impedances along the seismic 2d profile. evaluation of the absolute acoustic impedance inversion result is shown with the acoustic impedance log inserted at the well location (fig. 4). an excellent fit is seen between the inverted log trace (blue), the low frequency model log trace (green) and the well log trace (red) at the top (4.4 sec twt) and bottom (4.45 sec twt) of the target sandstone interval. however, it is important to notice that the inversion result (blue line) underestimates the absolute acoustic impedance in the deeper parts of the sandstone interval. this will lead to an overestimation of the porosity in this interval. distance (km) t w o -w ay t ra ve l ti m e (s ec ) 10 10.5 11 11.5 12 12.5 13 13.5 14 4.4 4.5 4.6 4.7 4.8 t o ta l p o ro si ty , p h it ( % ) 0 5 10 15 20 25 phit random line svane-1a area intra-kimmeridgian marker svane-1a wsw ene fig. 6. close-up of the 2d total porosity (phit) variation around the svane-1a well, showing the phit variation below the intra-kimmeridgian marker horizon along part of the inverted 2d seismic profile between faults a and c (figs 7, 8). the high porosity (>25%) between 4.55 and 4.6 sec twt at the well location is an artefact due to underestimation of the absolute acoustic impedance in the inversion (see text and fig. 4). the shape of the phit profile is governed by the limited depth interval 5300–5628 m, for which the applied ai-phit transform is defined, corresponding to a time window of 200 msec below the intra-kimmeridgian marker horizon. 2 km0 low seismic amplitude high amp. 1000 2000 3000 4000 5000 6000 bb aa ccb a svane-1asvane-1asvane-1a nnw–sse-striking high amplitude anomalies continue to the north of the svane-1 area c 04°24´36.22´04°24´36.22´ 56°07´50.56´56°07´50.56´ 04°24´36.22´́ e´e 56°07´50.56´́ n´n 04°24´36.22´́ e 56°07´50.56´́ n fig. 7. seismic amplitude extract from a 65 msec time window below the intra-kimmeridgian marker horizon. nnw–sse-trending high amplitude anomalies (blue colours) continue to the north of the svane-1a well. a, b and c mark the location of faults (see fig. 8). a b c 5.5 5.0 tw o -w ay t ra ve l ti m e (s ec ) 4.5 enewsw a svane-1asvane-1a b c 2500 m intra-kimmeridgianintra-kimmeridgian markermarker intra-kimmeridgian marker svane-1a fig. 8. close-up of the 2d seismic section across the svane-1a well (black). notice the high amplitude reflections below the intra-kimmeridgian marker horizon (yellow) between faults a, b and c (blue). 1616 calculating porosity from acoustic impedance cross-plotting the log-derived acoustic impedance (ai) and the log-derived total porosity (phit) using the svane-1a sonic and density log data from the depth interval 5300– 5628 m results in an ai-phit transform obtained from a second-order polynomial regression line (fig. 5): phit = 0.8176 – 1 × 10–4ai + 3 × 10–9ai2 where porosity is given in fraction and ai in g cm–3 × msec–1. the ai-phit transform, which is valid for the limited depth interval 5300–5628 m corresponding to a time window of 200 msec below the intra-kimmeridgian marker horizon, can be applied to convert the 2d acoustic impedance inversion result into a 2d total porosity (phit) profile in this time window (fig. 6). this 2d phit profile is a closeup of the central part of the inverted seismic profile at the 4-dip closure. the phit profile shows the lateral distribution of porosity anomalies below the intra-kimmeridgian marker horizon around the well location. alternating high (15–25%) and low porosity (5–15%) layers are seen below the intra-kimmeridgian marker horizon. in the time window 4.47–4.51 sec twt, the modelled porosities are up to 5% too high as a consequence of underestimation of the absolute acoustic impedance in the deeper parts. the derived total porosity values show good agreement with the observed svane-1a well porosities (up to 20–22% in the upper sandstone units), and indicate the presence of high porosity intervals off-structure down along the flanks of the structural high. porosity prediction tool the existence of upper jurassic sandstones with high porosities (15–25%) has been demonstrated in the svane-1a well and interpreted from the inversion result. thus an important question concerns the lateral extension and distribution of these sand-rich layers, and the challenge is to predict the location of yet undrilled high-porosity sandstone layers away from the well. for this purpose, a seismic amplitude extraction map for a narrow time window of 65 msec below the intra-kimmeridgian marker horizon was created to illustrate the lateral distribution of porosity anomalies in the vicinity of the svane-1 well (figs 7, 8). the seismic amplitude extraction map indicates that high amplitudes associated with the high porosity sandstone unit are concentrated along nnw–sse-trending anomalies that extend to the north of the svane-1a area (fig. 7). a close-up of the seismic data shows a correlation between fault planes (fig. 8 a–c) and low amplitude features on the amplitude extraction map (fig. 7). the amplitude extraction map also implies high lateral variability in the distribution of porosity anomalies corresponding to lateral variations in reservoir quality over the area. a possible new target area for further exploration could thus be located further to the north along the svane structure where high amplitudes prevail. the application of seismic inversion data based on well-log data, seismic data and a thorough geological model can significantly increase the possibility for finding new targets. references dolberg, d.m., helgesen, j., hanssen, t.h., magnus, i., saigal, g. & pedersen, b.k. 2000: porosity prediction from seismic inversion, lavrans field, halten terrace, norway. the leading edge 19, 392–399. japsen, p., britze, p. & andersen, c. 2003: upper jurassic – lower cretaceous of the danish central graben: structural framework and nomenclature. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 233–246. johannessen, p.n., dybkjær, k., andersen, c., kristensen, l., hovikoski, j. & vosgerau, h. 2010: upper jurassic reservoir sandstones in the danish central graben: new insights in distribution and depositional environments. in: vining, b.a. & pickering, s.c. (eds): petroleum geology: from mature basins to new frontiers. proceedings of the 7th petroleum geology conference, 127–143. london: geological society. doi: 10.1144/ 0070127. maver, k.g. & rasmussen, k.b. 1995: seismic inversion for reservoir delineation and description. society of petroleum engineers paper, 10 pp. doi: 10.2118/29798-ms. møller, j.j. & rasmussen, e.s. 2003: middle jurassic – early cretaceous rifting of the danish central graben. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 247–264. petersen, h.i., nytoft, h.p., vosgerau, h., andersen, c., bojesen-koefoed, j.a. & mathiesen, a. 2010: source rock quality and maturity and oil types in the nw danish central graben: implications for petroleum prospectivity evaluation in an upper jurassic sandstone play area. in: vining, b.a. & pickering, s.c. (eds): petroleum geology: from mature basins to new frontiers. proceedings of the 7th petroleum geology conference, 95–111. london: geological society. doi: 10.1144/ 0070095. rasmussen, k.b. & maver, k.g. 1996: direct inversion for porosity of post stack seismic data. society of petroleum engineers paper, 12 pp. doi: 10.2118/35509-ms. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: tab@geus.dk geological survey of denmark and greenland bulletin 26, 2012, 17-20 17 rock-cored drumlins on bornholm, denmark peter roll jakobsen the surface morphology of denmark is predominantly of glacial origin, created in depositional, deformational and erosional environments. in addition, postglacial marine, freshwater and aeolian processes have formed a variety of landforms. overviews of the danish landscape were published as geomorphological maps (milthers 1948; schou 1949; smed 1981), and a new one is currently in preparation. on bornholm, the morphology differs from the rest of the country because bedrock is present at or near the surface. this paper describes drumlins formed on bedrock on bornholm, which have not previously been recognised. new geomorphological map of denmark 1:200 000 the purpose is to create a map based on the geographical information systems (gis). to draw the boundaries of the different landform types as precisely as possible, the morphology was re-interpreted on the basis of topographical maps, lidar (light detection and ranging) data, geological maps as well as available literature. the scale of the new map is 1:200 000, but it is compiled at 1:100 000 and will be published in both a printed and a digital version, the former in four map sheets (jacobsen in press): (1) northern jylland, (2) central jylland, (3) southern jylland and fyn and (4) sjælland, surrounding islands and bornholm. so far, a preliminary version of southern jylland has been published (gravesen et al. 2004), and the map sheet covering sjælland, the surrounding islands and bornholm is completed. most of the surface features have already been described (milthers 1948; schou 1949; smed 1981). however, the morphological elements shown on these maps vary to some degree. the new geomorphological map of denmark also differs significantly from older maps, and some surface elements are re-interpreted. new elements include mega-scale glacial lineations and rogen moraines, and on lolland a valley formerly classified as a tunnel valley is re-interpreted as a fracture valley. bornholm large parts of bornholm consist of pre-quaternary crystalline bedrock with a discontinuous, thin cover of quaternary sediments and is classified as glacially scoured bedrock on the geomorphological map (fig. 1). during the quaternary glaciations, the bedrock was affected by overriding glaciers. it mainly shows erosional features such as fracture valleys and mega-scale glacial lineations, but also smaller features, not shown on the map, such as roches moutonnées and glacial striae (grönwall & milthers 1916; gravesen 1996). one of the most pronounced terrain features on bornholm is fracture valleys, which are subglacially eroded faults and fracture zones within the basement rocks. these valleys outline the fault systems from multiple deformation phases in the sorgenfrei–tornquist zone (grönwall & milthers 1916; graversen 2009). the main orientations of the fracture valleys are n–s, nne–ssw and ne–sw. a few are oriented nw–se. on bornholm, the orientations of glacial striae show two general directions: ne–sw and ese–wnw (fig. 2; grönwall & milthers 1916). mega-scale glacial lineations within the glacially scoured bedrock terrain are parallel with the measured glacial striae, and most of them are oriented ne–sw. in general, the two populations of glacial striae on bornholm occur in two different groups of pre-quaternary rocks. in the southern part of bornholm, the pre-quaternary geology is dominated by sandstone, shale and unconsolidated or poorly consolidated sediments, which are softer than the granites and gneisses in the rest of the island. the southern part of bornholm is dominated by a till plain. mega-lineations in this till plain show a slightly curved ese–wnw trend and are parallel to the dominant orientation of glacial striae in the area (figs 1, 2). marginal moraines are not as distinct as in other parts of denmark. they occur as scattered sandy and gravelly hills, and outline three ice marginal stages in the central and northern parts of bornholm (fig. 2). a n–s-trending ridge in the south-eastern part of bornholm, associated with a kame, is also interpreted as a marginal moraine (fig. 1). it is a low ridge that separates a till plain with mega-lineations to the west from a smooth till plain to the east. a few small outwash plains are found in south-western bornholm. raised beaches and beach ridges from the baltic ice lake are found along the coast at different levels, especially towards the east and north (fig. 1). © 2012 geus. geological survey of denmark and greenland bulletin 26, 17–20. open access: www.geus.dk/publications/bull 1818 rock-cored drumlins in the northernmost part of bornholm, within the glacially scoured bedrock terrain, there are ten elongated hills with typical drumlin shapes (figs 1, 3). drumlins are subglacial bedforms that were generated by the activity of overriding ice. drumlins are oval-shaped hills that formed beneath an ice sheet and aligned in the direction of ice flow, and they are common and widespread in formerly glaciated regions. the formation of drumlins has been widely discussed, and they are one of the most studied glacial landforms on earth (clark et al. 2009; johnson et al. 2010). on bornholm, lidar data have been invaluable in recognising these features (fig. 3). like other drumlins, those on bornholm are oval, and the surface has smooth contours with gradual fall to all mega-scale glacial lineation till plain drumlin esker kame marginal moraine outwash plain valley system glacial lake plain raised beach plain (baltic ice lake) raised beach ridge (baltic ice lake) beach ridge marine plain bog dune aeolian plain township fracture valley glacially scoured bedrock 5 km 10° 14°e 56° 57°n sweden germany 50 km jylland fyn bornholm 55°15´n 55°15´n 14°58´e 14°58´e denmark sjælland fig. 1. geomorphological map of bornholm. 19 sides. the length of the drumlins varies from 140 m to a little more than 600 m and their heights from 5 m to 31 m (table 1). their length to width ratio varies from 1.6 to 2.9 with an average of 2.3. the orientation of the hills is ne–sw, with an average of 50°. the drumlins have a core of granite and are more or less covered by till (fig. 3) that has predominantly been deposited towards the south-west on the lee side of the drumlins. the ice movement direction, indicated by glacial striae, is from the north-east (grönwall & milthers 1916), and the till cover is thickest on the south-western lee side of the drumlins. discussion in the northern part of bornholm, the drumlins and glacial striae indicate an ice-movement direction from the north-east to the south-west and are developed on crystalline bedrock. a few glacial striations have an e–w orientation (fig. 2). in the southern part of bornholm, only ice-movement directions from e to w are seen within the softer bedrock. these e–w erosional features were formed during a subsequent ice-stream event, which erased earlier ne–sw features in the southern part, but did not have the capacity to erode the harder crystalline bedrock. it appears that the subglacial conditions during the ice advances from the north-east were suitable for erosion of the granites on the northern part of bornholm. as indicated by the fracture valleys, the fault pattern is in n–s, nne–ssw and ne–sw directions in this part of bornholm, which is probably also true for the joint pattern. this joint and fault pattern could also favour formation of the elongated hills in this ne–sw direction, and would result in a preferred direction of erosional features by the ice advances from the north-east. the orientations of the erosional features range from 39° to 65° and could reflect several ice advances, although they cannot clearly be separated into two or more directions. investigations of drumlins in sweden (hättestrand et al. 2004) show that rock-cored drumlins are formed by successive phases of erosion. if this also holds for the drumlins of bornholm, then they may have been formed during several ice advances or even ice ages, and only by glaciers advancing from the north-east. conclusions on the northern part of bornholm, ten bedrock-cored drumlins are recognised and are included in the new digital, geomorphological map of denmark at a 1:200  000 scale. the lidar data were of great value in recognising them, although they are also recognisable from the contours of the topographic maps. the drumlins have a core of granite and are more or less covered with till that was predominantly deposited towards the south-west on the lee side of the drumlins. the lengths vary from 140 m to 612 m, and the height from 5 m to 31 m. the orientation of the drumlins is ne–sw, with an average strike of 50° formed by subglacial processes during one or several ice advances from the north-east. the orientation of the fracture valleys could very well intensify the preferred erosional orientation parallel to the drumlins, and thus be an important factor in the drumlin formation. ice marginal stage granite and gneiss sandstone, shale and sediments glacial striae pre-quaternary geology 50 km n fig. 2. map of bornholm showing orientation of glacial striae on bedrock, the distribution of crystalline and sedimentary bedrock types and ice marginal stages (after grönwall & milthers 1916; hansen & poulsen 1977). number/ length width height l:w orientation name (m) (m) (m) ratio (°) 1 langebjerg 612 211 31 2.9 60 2 kælderbakke 140 67 7 2.1 65 3 høje meder 177 102 25 1.7 46 4 kajbjerg 332 160 10 2.1 48 5 byggehøj 532 205 8 2.6 46 6 hestenshøj 330 145 10 2.3 56 7 blåholtshus 395 137 10 2.9 51 8 brogård 225 98 5 2.3 47 9 blåholtsgård 391 148 10 2.6 43 10 hammershus 389 239 25 1.6 39 table 1. dimensions of the drumlins on bornholm 2020 references clark, c.d., hughes, a.l.c., greenwood, s.l., spagnolo, m. & ng, f.s.l. 2009: size and shape characteristics of drumlins, derived from a large sample, and associated scaling laws. quaternary science reviews 28, 677–692. gravesen, p. 1996: geologisk set – bornholm. en beskrivelse af områder af national geologisk interesse, 208 pp. københavn: miljøministeriet, skovog naturstyrelsen og danmarks og grønlands geologiske undersøgelse. gravesen p., jakobsen, p.r., binderup, m. & rasmussen, e.s. 2004: geologisk set – det sydlige jylland. en beskrivelse af områder af national geologisk interesse, 188 pp. københavn: miljøministeriet, skovog naturstyrelsen og danmarks og grønlands geologiske undersøgelse. graversen, o. 2009: structural analysis of superimposed fault systems of the bornholm horst block, tornquist zone, denmark. bulletin of the geological society of denmark 57, 25–49. grönwall, k.a. & milthers, v. 1916: beskrivelse til geologisk kort over danmark i maalestok 1:100  000, kortbladet bornholm. danmarks geologiske undersøgelse i. række 13, 281 pp. + atlas. hansen, m. & poulsen, v. (eds) 1977: geologi på bornholm, 96 pp. varv ekskursionsfører 1. københavn: tidsskriftet varv. hättestrand, c., götz, s., näslund, j.o., fabel, d. & stroeven, a.p. 2004: drumlin formation time: evidence from northern and central sweden. geografiska annaler 86a, 155–167. jakobsen, p.r. in press: geomorphological map of denmark, 1:200 000, 4 map sheets. copenhagen: geological survey of denmark and greenland. johnson, m.d., schomacker, a., benediktsson, í.ö., geiger, a.j., ferguson, a. & ingólfsson, ó. 2010: active drumlin field revealed at the margin of múlajökull, iceland: a surge-type glacier. geology 38, 943–946. milthers, v. 1948: det danske istidslandskabs terrænformer og deres opstaaen. danmarks geologiske undersøgelse iii. række 28, 233 pp. schou, a. 1949: atlas of denmark, i, landscapes, 160 pp. + atlas. copenhagen: det kongelige danske geografiske selskab. smed, p. 1981: landskabskort over danmark, sheet 4, sjælland, lolland, falster, bornholm. brenderup: geografforlaget. 1 2 3 4 5 6 7 8 9 10 a b granite clay till meltwater sand outwash sand peat aeolian sand beach deposits from the baltic ice lake 500 m 500 m n n fig. 3. a: lidar map of the northernmost part of bornholm. the elevation varies from sea level to 118 m, and the highest parts are shown in white. the drumlins are indicated with red lines. the numbers refer to table 1. b: geological map of the same area. author’s address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: prj@geus.dk geological survey of denmark and greenland bulletin 29, 2013, 124 pp. 1 geological survey of denmark and greenland bulletin 29 • 2013 tectono-magmatic evolution of the younger gardar southern rift, south greenland brian g.j. upton geological survey of denmark and greenland danish ministry of climate, energy and building 22 geological survey of denmark and greenland bulletin 29 keywords troctolite, nepheline syenite, quartz syenite, alkali granite, agpaite, continental rifting, cumulates, mesoproterozoic. cover illustration view east-north-east from the tugtutôq central complex to the ilímaussaq and narssaq complexes. the igdlerfigssalik complex is visible in the far right distance. the lake (store pilesø) stretching away from the viewer overlies a sector of the older giant dyke. frontispiece: facing page view towards the igdlerfigssalik complex with fresh autumn snow from kongevejen near igaliku, displaying its characteristic, dull grey appearance and magmatic layering. photo:a.a. garde. chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, department of earth sciences, university of aarhus; minik rosing, geological museum, university of copenhagen; finn surlyk, department of geosciences and natural resource management, university of copenhagen scientific editors of this volume: lotte m. larsen and adam a. garde editorial secretaries: jane holst and esben w. glendal referees: john c. bailey (dk) and tom andersen (no) illustrations: eva melskens digital photographic work: benny m. schark layout and graphic production: kristian a. rasmussen printers: rosendahls-schultz grafisk a/s, albertslund, denmark manuscript received: 18 july 2012 final version approved: 1 july 2013 printed: 8 november 2013 issn 1604-8156 isbn 978-87-7871-366-7 citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 29, 124 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk and at www.geus.dk/publications/bull © de nationale geologiske undersøgelser for danmark og grønland (geus), 2013 for the full text of the geus copyright clause, please refer to www.geus.dk/publications/bull mailto:geus@geus.dk http://www.geus.dk/bull http://www.geus.dk/bull 3 44 contents abstract . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7 introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9 uniqueness of the southern branch of the gardar rift . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9 nomenclature of place names . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10 general geological overview . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10 history of exploration . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 14 gravity map . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 14 the older giant dyke complex, tuttutooq . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 15 marginal facies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16 central facies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16 the younger giant dyke complex . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19 composition of the magma . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 20 crystallisation sequence . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 21 internal structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23 the ygdc in the tuttutooq archipelago . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23 sissarluttooq . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 24 marraat . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25 asorutit . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25 krydssø . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 26 itillip saqqaa . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 28 tripyramidal peak west-south-west of itillip saqqaa . . . . . . . . . . . . . . . . . . . . . . . . . . . 33 itillinnuujuk . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 33 minor offshoots from the giant dykes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 34 narsaq gabbro and lopolithic relicts . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 34 younger giant dyke extensions west and north of motzfeldt sø . . . . . . . . . . . . . . . . . . . . . . 37 sydtungegletscher and syenitknold . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 38 central complexes and late dykes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 41 klokken complex . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 41 marginal gabbro . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 41 syenogabbro and unlaminated syenite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 41 central layered series . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 43 anorthosite xenoliths and plagioclase megacrysts in the ygdc and klokken gabbros . . . 46 origin of synformal layering in the younger giant dyke complex . . . . . . . . . . . . . . . . . . . . 47 mela-aillikites, carbonate-silicate rocks and carbonatites . . . . . . . . . . . . . . . . . . . . . . . . . . . 48 mela-aillikite intrusions in the narsaq area and on tuttutooq . . . . . . . . . . . . . . . . . . . . . . 48 mantle xenoliths . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 50 diatremes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 50 other aillikite, carbonate-silicate and carbonatite dykes . . . . . . . . . . . . . . . . . . . . . . . . 51 genesis of the ultramafic rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 52 narssaq complex . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 52 south qôroq complex . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 54 post-ygdc dyke swarms . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57 5 main dyke swarm . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57 big feldspar dykes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 58 salic dykes of the main dyke swarm . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 61 igaliko dyke swarm . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 62 tugtutôq central complex . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 64 mineralogy and geochemistry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 67 petrogenesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 68 late basic dykes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 69 ilímaussaq complex . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 70 augite syenite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 73 alkali granite and quartz syenite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 74 agpaitic syenites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 75 roof series . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 75 floor series . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 79 hyperagpaites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 87 hidden layered series . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 88 ilímaussaq parental magma . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 89 micro-kakortokite dyke . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 89 østfjordsdal syenite and igdlerfigssalik complex . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 90 age relationships . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 90 østfjordsdal syenite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 90 igdlerfigssalik complex . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 90 the role of anorthosite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 94 laminated anorthosites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 95 genesis of the anorthosites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 96 emplacement mechanisms and tectonics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 98 emplacement mechanism of the giant dykes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 98 tectonics within the southern rift . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 100 normal faulting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 101 transcurrent faulting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 101 evolution of the magmatic system of the younger gardar southern rift . . . . . . . . . . . . . . 103 parental mafic magmas . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 103 geochemical characteristics of the southern rift mafic magmas . . . . . . . . . . . . . . . . . . . . 103 magma evolution in the southern rift zone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 104 magmatic differentiation in the lower crust . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 106 magma chambers of the central complexes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 107 genesis of the silica-oversaturated magmas . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 108 crystallisation histories . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 108 mantle sources . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 109 rifting of the columbia supercontinent . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 111 topography of the younger gardar southern rift . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 111 summary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 114 acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 115 references . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 116 66 7 abstract upton, b.g.j. 2013: tectono-magmatic evolution of the younger gardar southern rift, south greenland. geological survey of denmark and greenland bulletin 29, 124 pp. the 1300–1140 ma gardar period in south greenland involved continental rifting, sedimentation and alkaline magmatism. the latest magmatism was located along two parallel rift zones, isortoq– nunarsuit in the north and the tuttutooq–ilimmaasaq–narsarsuaq zone in the south addressed here. the intrusive rocks crystallised at a depth of <4 km and are essentially undisturbed by later events. magmatism in the southern zone began with the emplacement of two giant, ≤800 m wide dykes and involved intrusion of transitional olivine basaltic, high al/ca magmas crystallising to troctolitic gabbros. these relatively reduced magmas evolved through marked iron enrichment to alkaline salic differentiates. in the older giant dyke complex, undersaturated augite syenites grade into sodalite foyaite. the larger, c. 1163 ma younger giant dyke complex (ygdc) mainly consists of structureless troctolite with localised developments of layered cumulates. a layered pluton (klokken) is considered to be coeval and presumably comagmatic with the ygdc. at the unconformity between the ketilidian basement and gardar rift deposits, the ygdc expanded into a gabbroic lopolith. its magma may represent a sample from a great, underplated mafic magma reservoir, parental to all the salic alkaline rocks in the southern rift. the bulk of these are silica undersaturated; oversaturated differentiates are probably products of combined fractional crystallisation and crustal assimilation. a major dyke swarm 1–15 km broad was intruded during declining crustal extension, with decreasing dyke widths and increasing differentiation over time. intersection of the dyke swarm and e–w-trending sinistral faults controlled the emplacement of at least three central complexes (narssaq, south qôroq and early igdlerfigssalik). three post-extensional complexes (tugtutôq, ilímaussaq and late igdlerfigssalik) along the former rift mark the end of magmatism at c. 1140 ma. the latter two complexes have oblate plans reflecting ductile, fault-related strain. the tugtutôq complex comprises quartz syenites and alkali granites. the ilímaussaq complex mainly consists of nepheline syenite crystallised from highly reduced, fe-rich phonolitic peralkaline (agpaitic) magma, and resulted in rocks with very high incompatible element concentrations. abundant anorthositic xenoliths in the mafic and intermediate intrusions point to a large anorthosite protolith at depth which is considered of critical importance in the petrogenesis of the salic rocks. small intrusions of aillikite and carbonatite may represent remobilised mantle metasomites. the petrological similarity between older and younger gardar suites implies strong lithospheric control of their petrogenesis. the parental magmas are inferred to have been derived from restitic ketilidian lithospheric mantle, metasomatised by melts from subducting ketilidian oceanic crust and by small-scale melt fractions associated with gardar rifting. there are numerous analogies between the southern gardar rift and the palaeogene east african rift. author’s address university of edinburgh, school of geosciences, grant institute, the king’s buildings, west mains road, edinburgh eh9 3jw, uk. e-mail: brian.upton@ed.ac.uk mailto:brian.upton@ed.ac.uk 88 fi g. 1 . o ve rv ie w m ap sh ow in g i m po rt an t p la ce n am es (g iv en in ’n ew sp el lin g’ ) u se d in th e t ex t, an d lo ca tio ns o f d et ai l m ap s w ith fi gu re n um be rs . 49 67 10 71 11 5 11 4 5 23 81 99 10 8 17 37 65 motz fel dt sø tu ttu to oq br ed efj ord se rm ilik n un ar su it n ar sa rs ua q g .f. h ol m n un at ak iv itt uu t q as sim iu t d yr næ s kr in gle rn e n ar sa q q aq or to q qo or qq up se rm ia qooroq sk ov fjo rd iso rto q ba ng h av n melle mlandet tre tu ng eg let sc he r g eo lo gf je ld k va ne fje ld ka ng erl ua rsu k in la nd ic e iga liku f jord tu nu llia rfik 61 ° 61 ° 47 ° n 9 introduction uniqueness of the southern branch of the gardar rift the concept of ‘the gardar rift’ is sometimes spoken of. no such single ‘gardar rift’ exists, but continental rifting certainly affected part of what is now southern greenland. the rifting affected a stretch of the crust that embraced the southern margins of the archaean craton and the adjacent regions of younger rocks that lie to its south. in the mesoproterozoic, which encompasses the gardar period between about 1300 and 1100 million years ago, greenland was a component part of the columbia supercontinent (rogers & santosh 2002) that lasted from c. 1800 million years (ma) until c.1100 ma. the affected crust was clearly a zone of mechanical weakness, vulnerable to repeated fracturing in response to stresses induced by the slow but steady flowage of hot mantle rocks deep beneath it. crustal fracturing and pull-apart caused the pressure on the underlying mantle rocks (peridotites) to be locally reduced, with the result that the mantle underwent partial melting. thus rifting was inevitably accompanied by magmatism which, in turn, caused volcanic activity at the surface. these processes were fundamentally the same as those operating within the paleogene east african rift system. this bulletin attempts to describe and interpret the faulting and associated magmatism that defined the more southerly of the two rift zones, generated in the later part of the gardar history, between 1180 and 1140 ma. what we can now study in the field are rocks that then lay at estimated depths of 3–4 km below their contemporary land surface but which were subsequently uplifted and eroded to their present positions. to the question “what makes these riftzones so interesting?” there are several answers. first of all, the late gardar southern rift may be globally unique in that nowhere else has an ancient rift been so dissected by erosion to reveal its deep anatomy. this alone accords it inestimable scientific interest. furthermore, it may be inferred that immense volumes of magma were generated, the greater part of which was retained deep in the crust to undergo slow cooling and chemical maturation. most of the latter took place via the process of fractional crystallisation by which the component elements (virtually all of the 92 elements that occur in nature) underwent selective redistribution. some of the latest melt fractions within the ilímaussaq complex crystallised to rocks of extreme compositions. the latter contain high concentrations of many of the planet’s rarest elements, including uranium, thorium, tin, niobium, beryllium, zirconium and the so-called ‘rare-earth elements’. consequently ilímaussaq has exercised a strong attraction for, not only petrologists, mineralogists and geochemists, but also mining prospectors and engineers. the igneous intrusions of the younger gardar southern rift were studied by the geological survey of greenland during its regional 1:20  000 geological mapping programme, 1956–1962. the mapping was followed by a large number of publications and unpublished phd theses. a hundred years have now passed since the publication of ussing’s very perceptive memoir on ilímaussaq (ussing 1912), and the current economic interest in south greenland offers an appropriate opportunity to collate the information gathered on this remarkable younger gardar southern rift. this has a generalised width of between 10 and 15 km and transects the ketilidian granitic julianehåb batholith which it post-dates by c. 700 ma (figs 1, 2). of the ultimate causes of the lithospheric extension that marked the early stages of rifting and the left-lateral faulting that essentially ended it, we remain ignorant. this memoir merely describes and interprets the attendant tectono-magmatic phenomena while leaving these fundamental questions unresolved. the growth of the whole great volcanic system, which undoubtedly spanned many millions of years, was followed by over 1100 ma of nearly unbroken quiescence. this extraordinary region has a character that should allow it to be granted world heritage status as a ‘geopark’. quite apart from all it offers, scientifically and potentially commercially, it is undoubtedly a region of great natural beauty, as yet virtually unspoiled. 1010 nomenclature of place names throughout this bulletin place names are written with modern greenlandic spelling. most of the gardar igneous complexes were named before the spelling reform in 1973, and because of the conventions for naming geological units the names of these geological features have not been changed. thus, ilimmaasaq denotes the mountain whereas ilímaussaq is the intrusion. accordingly, referring to the geographical features, the rift system described here will appear as the tuttutooq–ilimmaasaq– narsarsuaq system. general geological overview wegmann (1938) divided the geology of southern greenland into ‘an old basement’ and younger formations. the former comprises archaean gneisses whilst wegmann subdivided the latter into the ketilides and the gardar formation. the palaeoproterozoic ketilidian rocks (c. 1800 ma) are bounded to the north by the archaean (>2800 ma) craton. emplacement of the ‘andean type’ julianehåb batholith marked the peak of ketilidian orogenic activity (garde et al. 2002). the batholith growth commenced at c. 1855 ma and was finished at c. 1795 ma, succeeded by fore-arc deposition, deformation, metamorphism and emplacement of the 1755–1723 ma rapakivi suite (garde et al. 2002). over the next c. 500 ma, the resultant co500 km ice ketilidian orogen metasedimentary rocks, pelite zone metasedimentary rocks, psammite zone julianehåb batholith rapakivi granite suite supracrustal rocks, border zone nunarssuit bangs havnbangs havn ilímaussaq narssaq tugtutôq south paatusoqklokken igdlerfigssalik orthogneiss–amphibolite complex archaean older gardar intrusions and eriksfjord formation gardar province isortoq–nunarsuit and tuttutooq– ilimmaasaq–narsarsuaq zones principal younger gardar complexes 60°n 61°n 62°n 100 km 46°w48°w 50°w 44°w 42°w kobberminebugt paatusoq nanortalik qaqortoq foreland border zone julianehåb batholith psammite zone pelite zone niaqornaarsuk kap farvel 61°n qôroqqôroq fig. 2. general geological map of south greenland with the mesoproterozoic gardar province within the palaeoproterozoic ketilidian orogen. note the isortoq–nunarsuit and tuttutooq–ilimmaasaq–narsarsuaq magmatic zones (rifts) comprising the labelled younger gardar complexes and the approximate extent of associated dyke swarms. the gardar intrusions have mainly been emplaced into the ketilidian julianehåb batholith. the ketilidian border zone comprises clastic and chemical sedimentary rocks deposited on an archaean basement, overlain in the west by a thrust sheet of metabasaltic and related rocks. modified from garde et al. (2002, fig. 1). 11 lumbia supercontinent experienced equilibration, uplift and erosion culminating in a long sequence of rifting events involving faulting and magmatism. the latter gave rise to the gardar igneous province in which suites of (mostly) genetically related alkaline rocks were emplaced. the spatial relationship of the mesoproterozoic gardar alkaline rocks to the palaeoproterozoic and archaean formations are shown in fig. 2 and radiometric age data are given in table 1. the earliest rocks that were ascribed to the gardar period by wegmann (1938) are the terrestrial sandstones and lavas composing the eriksfjord formation that unconformably overlie the julianehåb batholith. isotopic dating of the lavas by paslick et al. (1993), using the smnd mineral and whole-rock method, gave ages of c. 1170 ± 30 ma and 1120 ± 30 ma. the qassiarsuk carbonatite–alkaline silicate volcanic complex, correlated with the mussartût lavas close to the base of the eriksfjord formation, is dated at c. 1200 ma (rb-sr and pb-pb; andersen 1997). the formation is, however, clearly cut by the motzfeldt pluton which has yielded ages of 1273 ± 6 ma (u-pb zircon; mccreath et al. 2012), 1282 ± 30 ma (blaxland et al. 1978) and 1226 ± 27 ma (finch et al. 2001b) (both rb-sr mineral & whole-rock analyses). u-pb zircon, baddeleyite and pb-pb pyrochlore age data paatusoq syenite pluton zircon 1144.1 1.1 m. hamilton (unp.) østfjordsdal syenite pegmatite zircon 1147.5 3.2 salmon (2013) tugtutôq granite pluton zircon 1156 1.1 l. heaman (unp.) ilímaussaq agpaite cumulate baddeleyite 1160 5 waight et al. (2002) ilímaussaq agpaite cumulate baddeleyite 1160 2.3 krumrei et al. (2007) tugtutôq younger giant dyke baddeleyite 1163 2 m. hamilton (unp.) tugtutôq younger giant dyke baddeleyite 1165.7 1.2 l. heaman (unp.) klokken syenite zircon 1166 3 harper (1988) nunarssuit syenite pegmatite zircon 1171 5 finch et al. (2001a) tugtutôq older giant dyke baddeleyite 1184 5 l. heaman (unp.) north motzfeldt nepheline syenite zircon 1257.4 6.7 salmon (2013) motzfeldt nepheline syenite zircon, pyrochlore 1273 6 mccreath et al. (2012) kûngnât gabbro ring dyke baddeleyite 1275.2 1.8 l. heaman (unp.) motzfeldt syenite pegmatite zircon 1275.3 1.1 salmon (2013) tugtutôq bd dyke baddeleyite 1279 1.3 l. heaman (unp.) kangerluarsuk bd dyke baddeleyite 1280 3 l. heaman (unp.) qaqortoq bd dyke baddeleyite 1284 3 l. heaman (unp.) intrusion or locality rock unit material age (ma) ± (ma) reference rb-sr whole-rock and mineral age data† nunarssuit quartz syenites, granites§ whole-rock 1130 14 blaxland et al. (1978) ilímaussaq, late dyke, kvanefjeld monchiquite phlogopite 1134 17 larsen (2006) klokken gabbros, syenites whole-rock 1135 11 blaxland et al. (1978) igdlerfigssalik ‘late complex’ gabbros, syenites, nepheline syenites whole-rock 1142 15 blaxland et al. (1978) tugtutôq central complex quartz syenites, granites whole-rock 1143 36 blaxland et al. (1978) ilímaussaq agpaites, syenites, granites whole-rock 1143 21 blaxland et al. (1978) tugtutôq older giant dyke gabbros, syenites, nepheline whole-rock 1150 9 blaxland et al. (1978) south qôroq syenites, nepheline syenites whole-rock 1160 8 blaxland et al. (1978) ilímaussaq agpaite alkali feldspar, eudialyte 1160 2.3 waight et al. (2002) bangs havn gabbros, syenites whole-rock 1185 22 engell & pedersen (1974) qassiarsuk trachytes, carbonatites whole-rock, min. sep. 1205 12 andersen (1997) kûngnât gabbro, syenites whole-rock 1219 16 blaxland et al. (1978) ivigtût alkali granite, cryolite body whole-rock 1222 24 blaxland et al. (1978) north motzfeldt nepheline syenites whole-rock, min. sep. 1226 27 finch et al. (2001) ivittuut region lamprophyres, dolerites biotite 1250 18 patchett et al. (1978) north qôroq nepheline syenites, lujavrite biotite 1268 60 blaxland et al. (1978) early motzfeldt complex gabbro, nepheline syenite, lujavrite biotite 1282 30 blaxland et al. (1978) grønnedal–íka nepheline syenites, carbonatite biotite 1299 17 blaxland et al. (1978) table 1. radiometric age determinations from gardar igneous rocks errors are quoted at the 2 level. † rb-sr data are presented as simple regressions, recalculated for the decay constant of steiger & jäger (1977) § l. heaman provided regressions for four units of the nunarssuit complex. they all lie within error of the value for the nunarssuit syenite quoted here. unp.: unpublished data; min. sep.: mineral separates. 1212 it is also cut by the north qôroq pluton (rb-sr age 1268 ± 60 ma; blaxland et al. 1978) (table 1). accordingly whilst parts of the eriksfjord formation appear to be older than 1270 ma, the principal outcrop farther to the west-south-west may be substantially younger (paslick et al. 1993) although it predates the younger gardar dyke complex and the narssaq and ilímaussaq complexes. other gardar intrusions that give relatively old ages include the grønnedal–íka complex (1299 ± 17 ma, rb/ sr dating; blaxland et al. 1978) and three early alkali olivine dolerite dykes (the bd0 dykes of the geological survey of greenland) which yield u-pb baddeleyite dates of 1284 ± 3 ma, 1280 ± 3 ma and 1279 ± 1.3 ma (table 1). although neither grønnedal–íka nor the bd0 dykes have contact relationships with the eriksfjord formation, the fact that each is believed to be related to alkaline magmatism in rifting environments lends support to the concept that the eriksfjord formation may be distinctly older than the dates indicated by the sm-nd method (table 1; fig. 3) the julianehåb batholith has a pronounced syn-magmatic foliation that was exploited in younger gardar time by shearing and dyke intrusion, both along the northern margin of the batholith in the nunarsuit–isortoq region (bounded in the north by the kobberminebugt and and in the south by sermilik fjord), and more centrally in the batholith, embracing the tuttutooq archipelago, the ilimmaasaq–qassiarsuk peninsula, the narsarsuaq area and the nunataks north of motzfeldt sø. the latter region houses the southern rift with the tuttutooq–ilimmaasaq–narsarsuaq magmatic system (fig. 2). most of the late gardar intrusions lie within these two eneoriented zones and the two jointly compose a dissected, asymmetric rift zone, approximately 70 km wide. whilst most of the intrusions in the southern rift (tuttutooq– ilimmaasaq–narsarsuaq) are concentrated within a narrow zone (c. 15 km wide), its northern neighbour (nunarsuit–isortoq) is over twice as wide. isotopic evidence suggests that archaean crust underlies both rift zones (halama et al. 2004; krumrei et al. 2006). although the northern rift possesses great intrinsic interest, description of it is beyond the scope of this bulletin. there is very considerable geographic overlap between the igneous rocks of the younger gardar intrusions and those of the older gardar (fig. 2). the latter date principally from 1280 to 1250 ma (table 1; fig. 2) and are also not considered in any detail in this publication. that the julianehåb batholith is bounded to north and south by ketilidian metasedimentary and volcanic formations (fig. 2) suggests the possibility that it marked the site of a pre-gardar dome. it is postulated that lithospheric weakness beneath the two gardar rift zones permitted attenuation, fissuring and admission of mantle melts, producing parallel volcanic grabens from which several kilometres of cover have been removed by erosion (upton & blundell 1978; fig. 1a). the crustal extension across each of the two zones, indicated by their dyke swarms, was approximately 1.5 km. hence the total dilation across c. 70 km was c. 3 km or c. 4.3%. a large positive bouguer gravity anomaly, attributed to gabbroic rocks at shallow depths, characterises the southern rift system (blundell 1978; fig. 4), although no such gravity high is known from the nunarsuit–isortoq zone. comparable gravity highs are associated with the east african rift system and the oslo rift. the lack of an anomaly beneath the nunarsuit–isortoq zone is explicable if it was a more diffuse zone of attenuation, preventing large bodies of mafic magma from attaining shallow levels (upton & blundell 1978). in iceland, fissure eruption is commonly a prelude to increasing localisation of magma ascent and generation of a central-type volcano. it was recognised that these features represent consecutive parts of a single ‘magmatic system’, a concept elaborated by walker (1993). by analogy with these icelandic phenomena, the whole suite of intrusions along the younger gardar southern rift (c. 1160–1140 ma) is interpreted as a coherent, large-scale magmatic system. this unified system marked a tectonomagmatic event in which strain energy release, after rising to a maximum, was followed by an extended period of relaxation. this bulletin is based on the contention that the ilímaussaq complex is among the youngest components of a great tectono-magmatic system that accompanied continental rifting. in its more mature stages, this system underwent gradual change from fissuring and dyke emplacement towards emplacement of stocks and ring-dykes as extensional stresses diminished. although the two gardar rift zones share many features in common, there is a clear petrological difference in that the absence of phonolitic/foyaitic and subordinate ultramafic lamprophyre/carbonatite rocks in the northern rift zone contrasts with their importance in the tuttutooq–ilimmaasaq–narsarsuaq zone to the south. two specific foci of lithospheric weakness in the southern rift were provided by the intersection of the batholithic foliation and a set of transcurrent sinistral faults trending wnw–ese to w–e. apart from radiometric dating and intersections among the larger intrusions, two features invaluable for establishing the chronology within the system are the main dyke swarm and the sinistral 13 faults. figuratively, the entire tuttutooq–ilimmaasaq– narsarsuaq magmatic system may be compared to a symphonic work in several movements. an abrupt introduction that rose to a crescendo in the first movement gave way to a relatively quiescent ending with emplacement of the central tugtutôq, ilímaussaq and igdlerfigssalik igneous complexes. accordingly two principal stages are recognisable. the first was characterised by extension and dyke intrusion, brought to an end by transcurrent faulting, whereas the second stage saw the passive ascent of mainly salic magmas. during this post-faulting stage further extension was negligible. the gardar province lay north of the grenville front and so avoided tectonic deformation in the grenvillian orogeny. since the mesoproterozoic, it has remained remarkably unscathed and the overall state of preservation of the gardar plutons is excellent. evidence for this stability is provided by e.g. miarolitic cavities in pegmatites of the (older gardar) kûngnât complex that appear never to have been occupied or mineralised since their initial crystallisation, and by delicate acicular aegirine crystals in pegmatite geodes in the ilímaussaq complex, highly susceptible to seismic damage, that remain unbroken. some feldspars in the klokken complex remain sufficiently unaltered to allow the oldest age dating by the ar-ar method concordant with u/pb dating (parsons et al. 1988). crustal uplift in the mesozoic/cenozoic was followed by extensive erosion by the pleistocene ice sheets. retreat of these over the past few thousand years has revealed the gardar igneous rocks as superlative ‘time-capsules’. among the approximately one dozen central-type plutons, the ilímaussaq complex is by far the best known because it hosts an extraordinary large-scale concentration of rare minerals and their component rare elements. the rare mineral assemblages are contained in highly peralkaline (agpaitic) rocks that were the youngest products of the complex. the most enriched part of the complex (kvanefjeld) is now believed to contain the world’s second largest deposit of rare-earth elements and its sixth largest uranium deposit (parsons 2012). research into ilímaussaq has currently resulted in over two-hundred and thirty scientific publications. in this bulletin an attempt is made to describe the various features of the tectono-magmatic system (intrusions and fault movements) in chronological order. this encounters some problems, e.g. faults can be re-activated on numerous occasions, some igneous units are undated or imprecisely dated, and the anorthosite body that from xenolith evidence inferentially underlies the entire province cannot be meaningfully discussed before description of the igneous rocks that host the xenoliths. nonetheless this is the broad outline followed here. 1320 ? ? 1300 1280 1260 1240 1220 1200 1180 1160 1140 1120 ma eriksfjord fm. igd le rfi gs sa lik ø st fjo rd sd al tu gt ut ôq in tr us io n so ut h q ôr oq kl ok ke n ilí m au ss aq tu gt ut ôq y g d n un ar ss ui t q as sia rs uk tu gt ut ôq o g d ilí m au ss aq la te d yk e iv igt ût n or th q ôr oq n or th m ot zfe ld t kû ng nâ t m ot zfe ld t g rø nn ed al– ík a br ow n dy ke s pa at us oq fig. 3. age distribution of rb-sr and u-pb age determinations of gardar intrusions from various sources (see table 1). ogd: older giant dykes. ygd: younger giant dykes. original drawing with courtesy of a. bartels. 1414 history of exploration the following notes are culled from sørensen (1967) “on the history of exploration of the ilímaussaq intrusion, south greenland” and supplementary notes in sørensen (2001). the earliest significant investigations of the mineralogy of southern greenland were those of k.l. giesecke in the course of an expedition from 1806 to 1813. during this time he made two visits (1806 and 1809) to the ilimmaasaq area (giesecke 1910). his resultant collections were intercepted at sea during the napoleonic wars by the british navy and landed at leith. however, most of his specimens eventually arrived in copenhagen and other european cities. amongst the new minerals described were arfvedsonite, eudialyte and sodalite. in the course of k.j.v. steenstrup’s expeditions (1874–1890) the ilimmaasaq area was visited in 1874, 1876 and 1877. steenstrup was accompanied by g. holm and a. kornerup (steenstrup 1910; steenstrup & kornerup 1881) and from their work in the julianehåb district, a map of the ilimmaasaq area was published in 1881. a collection of rocks and minerals was made by flink in 1883 and some of the new minerals were described by bøggild & winther (1899). the commission for the geological and geographical exploration of greenland was continued in 1903 by n.v. ussing (in the company of o.b. bøggild) who mapped the intrusions around the ilimmaasaq area and igaliku. he also investigated nunarsuit, grønnedal and ivittuut. following his revisit to the ilimmaasaq area in 1908, ussing published his seminal work on the area in 1912. he recognised the principal rock types in ilímaussaq as augite syenite, alkali granite, pulaskite, foyaite, sodalite foyaite and the poikilitic sodalite syenite that he was to name naujaite. he also described ‘banded eudialyte nepheline syenites’, to which he bestowed the name kakortokites, and finegrained nepheline syenites (lujavrites). for the exotic eudialyte-bearing syenites ussing coined the collective name ‘agpaites’. ussing also recorded the essexites, nordmarkites and alkali granites in the vicinity of narsaq. although there were some further studies of the ilimmaasaq area by s.g. gordon (published in 1924) and r. bøgvad who visited it in 1939 on behalf of the company kryolitselskabet øresund, virtually no serious geological investigations were made into the igneous rocks of the area until after world war ii when the geological survey of greenland (ggu) became established in 1946. nonetheless, c.e. wegmann made very significant geological advances in the region in 1936 when he established the basis for a chronology, introducing the term gardar period, named after the norse archbishopric of gardar established in what was then called eriksfjord (wegmann 1938). all the igneous activity described in this bulletin falls into his gardar period. after the war, reconnaisance work was undertaken by a. noe-nygaard, k. ellitsgaard-rasmussen, r. bøgvad and h. sørensen and, subsequently, j. bondam and h. pauly. in 1955, after the danish government decided to investigate the potential uranium resources of the ilímaussaq complex, a systematic geological mapping programme was initiated that was completed in 1962. recognition of the enetrending zone of faulting, dyke intrusion and emplacement of salic complexes around tunulliarfik was first mentioned by berthelsen & noe-nygaard (1965). in the southern rift zone, dykes of exceptional width (‘giant dykes’) reaching from 500 to 800 m width have few or no phanerozoic counterparts. they exhibit the remarkable localised ‘ballooning’ and enigmatic changes, which suggests that the stress fields, lithospheric thicknesses and/or the thermal state of the lithosphere were dissimilar in the mesoproterozoic and phanerozoic. the giant dykes were initiated in the tuttutooq–ilimmaasaq–narsarsuaq rift by the most primitive basaltic magmas in the gardar province. the largest of these intrusions is the younger giant dyke complex or ygdc. gravity map a bouguer gravity map of the eastern gardar province around narsaq and julianehåb (fig. 4), contoured at 50 gravity unit intervals, was produced by blundell (1978). all its values are negative relative to the international geodetic reference field. the contours are broadly parallel to the coast line and indicate a regional gradient with values from –200 gravity units near the coast to –700 gravity units close to the inland ice. superimposed on this regional pattern is a linear gravity high some 50 km long, centred on tuttutooq and tunulliarfik. fortunately, the trend of this high is almost perpendicular to the regional gradient so that the two anomalies can be distinguished; but unfortunately, the gravity survey did not extend far enough inland to effect complete separation. the gravity high, though not as great as e.g. the north american gravity high (chase & gilmer 1973) or that of the kenya rift (fairhead 1976), is clearly distinct from the regional field and cannot be accounted for by the exposed gardar intrusions. it was interpreted as due to an underlying mass of dense mate15 rial, deduced to be mostly at a depth of 3 to 5 km below the present surface. the anomaly broadens and decreases in magnitude westwards but reaches its highest structural levels in the vicinity of narsaq. beneath tuttutooq it may represent the merging at depth of the giant dykes. the axis of the gravity high lies 3–5 km south of, but parallel to, the giant dykes. the intrusion responsible for the high underlies tuttutooq at relatively shallow depths and is manifest at the surface as the narsaq gabbro. blundell (1978) concluded that it extends down to a depth somewhere between 10 and 40 km. depending on the model chosen it is between 10 and 25 km wide, the latter representing the full width of the rift. the intrusion was inferred to include gabbro as a major component although, in view of the studies of the ygdc, peridotite cumulates probably play a large role. the character of the linear gravity high is consistent with those of other continental rifts (e.g. the oslo graben; ramberg 1972), adding weight to the view that the tugtutôq complex is an eroded continental rift (upton & blundell 1978). it also accords with the interpretation of the north american ‘mid-continent high’ in terms of a rift structure in which gabbroic intrusion is the major cause (ocola & meyer 1973). in view of the congruence between the linear gravity high and the ygdc outcrops in tuttutooq, it may be anticipated that a recurrence of a gravity high would be found corresponding to the ygdc extensions up to the inland ice (discussed below) but, because of logistical problems, no data are available for that region. nonetheless, the field evidence for largescale emplacement of mafic magma early in the younger gardar episode together with the geophysical data clearly indicate that the giant dykes and their subsurface extensions reflect a very major magmatic event. the older giant dyke complex, tuttutooq the older giant dyke complex (ogdc) is a massive, parallel-sided dyke, 500–600 m wide, with an undulating course traceable for c. 20 km through the island of tuttutooq (fig. 5; upton 1962; 1964c; upton et al. 1985). whilst its chilled marginal facies show that it commenced with intrusion of hawaiitic magma, notably enriched in incompatible elements, the intrusion as a whole comprises a wide array of alkaline rocks. although intrusion of the younger giant dyke complex, described below, is regarded as the major tectono-magmatic event in the development of the southern rift, emplacement of its closely related predecessor, the ogdc, is taken as the event that heralded rifting. u-pb baddeleyite dating on the ogdc has yielded an age of 1184 ± 5 ma whereas rb-sr age determinations gave 1154 ±16 ma and 1150 ± 9 ma (table 1). the u-pb baddeleyite date is accepted as the best age for the ogdc, also because the clearly younger ygdc has a u-pb baddeleyite age close to 1163 nunarssuit complex inland ice ilímaussaq igdlerfigssalik bouguer anomaly (gravity units) –300 – –400 –400 – –500 –500 – –600 –600 – –700 –700 – –800 –800 – –900 25 km 46°30´ 61° n fig. 4. gravity map around the younger gardar tugtutôq, narssaq and ilímaussaq complexes, contoured at 100 gravity unit intervals. simplified from blundell (1978, fig. 2). 1616 ma (table 1). thus, the beginning of the southern rift will be taken as 1184 ma. marginal sheaths or ‘border groups’ up to 100 m wide, consisting of gabbroic to ferro-syenogabbroic rocks, enclose a 300 m wide axial zone of salic rocks that grade from augite syenite in its western parts, through pulaskite and foyaite to peralkaline sodalite foyaite in the easternmost outcrop. the salic rocks in the centre of the intrusion may represent a continuous sequence of cumulates (upton et al. 1996) although a previous interpretation had suggested that the syenites crystallised in situ from a compositionally stratified magma body (upton et al. 1985). cryptic and phase layering phenomena within the salic rocks suggest a ‘way up’ from west to east and it is inferred that the intrusion was tectonically tilted (during late gardar block-faulting?) about an axis normal to its length so that after uplift and erosion a deeper section is exposed at the western end than in the east (upton 1962). the estimated difference in structural (‘stratigraphic’) levels is 2 to 3 km. marginal facies the symmetrically graded marginal facies varies in width from c. 50 m to 100 m along most of the intrusion but appears to be absent over a short distance to the east of the cross-cutting tugtutôq central complex. the rustybrown weathering mafic rocks had greater erosional resistance than the central facies and consequently tend to form upstanding ledges on either side of the intrusion. contacts against the julianehåb granite are sharply defined and well chilled and are best exposed in the coastal section provided by a fjord (sejlfjord) that bisects it approximately at right angles (figs 5, 6). away from the contact zones the rocks are subophitic and layering features are absent or very weakly developed. the feldspars are strongly zoned plagioclases surrounded by perthitic calcic anorthoclase and grading to cryptoperthite outermost zones, lacking discernible exsolution features. within these syenogabbroic rocks the feldspars exhibit a continuum from an63 via potassic oligoclase (an22) and calcic anorthoclase into ca-nasanidine (upton et al. 1985), whilst the olivines show a range of fo53–16. figure 7 shows the compositional range of olivines and pyroxenes. transition from mafic marginal facies to the felsic central facies occurs over a width of 1–2 m and involves a complex mélange of the two rock types with the felsic rocks forming an irregular network pattern that may have originated from sidewall ascent of buoyant felsic magma interacting with incompletely crystallised syenogabbro (fig. 8). central facies throughout its westernmost 12 km the central facies consists of mesocratic augite syenite although outcrops are sparse. the augite syenite is, however, splendidly exposed on either side of sejlfjord. preferential glacial excavation of the syenites left a broad flat valley and its light colouration gave rise to the term ‘the white valley’ or ‘hviddal’. the augite syenites contain perthitic feldspars up to 10 mm in size, with turbid (altered) interstitial nephelines. the ferromagnesian minerals are fayalitic olivine (largely replaced by iddingsite) with a compositional range of fo10–4, and idiomorphic clinopyroxene zoned from pinkish-grey centres to pale green rims that are typically surrounded by amphibole (brown, zoning out to blue-green) reaction rims. titanomagnetite with biotite fig. 5. a: geological map of the giant dyke complexes on tuttutooq and relationship to the younger central intrusions. b: reconstructed form of the ogdc before the later intrusions and faulting. dark brown: augite syenite. lighter brown: pulaskite. yellow: foyaite. gradual transition from north-west to south-east. younger giant dyke complex tugtutôq central complex older giant dyke complex julianehåb batholith fault fault a b tuttu tooq sto r e pil esø sejlfjord 45°30´ 60°51´ 17 reaction fringes and fluor-apatite are minor components. the amphibole and biotite are probably subsolidus reaction products (powell 1978). east of the tugtutôq central complex, the ogdc is seen only in scattered outcrops along the coasts of store pilesø and the small islands within it. the syenites in this sector have experienced substantial hydrothermal alteration, with development of epidote. the nepheline content is markedly higher (15–20% modally) than in the western syenites, and farther east the nepheline becomes increasingly idiomorphic. analyses of separated feldspars show them to be essentially ca-free, close to the ab–or join between or40 and or55 and straddling the na-sanidine/sanidine fields (upton 1964c). feldspars from the most easterly outcrops are the most potassic whereas the host rocks become increasingly sodic. these compositions do not precisely lie on the extrapolated trend from the marginal zone, suggesting some discontinuity between the marginal series and the central series. barium contents in the feldspars peak at c. 1.6 wt% ba within the na–ca-anorthoclase field, corresponding to c. 2.8 mol% celsian. the ogdc presents a near-complete spectrum of rocks from gabbroic to foyaitic that are, inferentially, products of a suite of magmas that graded from the initial hawaiitic magma through ferro-mugearitic, nephelinebenmoreitic, nepheline-trachytic to peralkaline phonolitic. analyses of the central series show negative eu anomalies but these are absent in the marginal series, implying that any calcic feldspar fractionation commenced relatively late in the evolution of the suite (upton et al. 1985). as the rocks are apparently devoid of lamination or modal layering that might be ascribed to crystal settling, the question as to whether or not they should be regarded as a cumulate sequence is open to debate. the alternative hypothesis is that the sequence composing the central zone developed from a stratified magma chamber fig. 6. view east across sejlfjord showing a section across the southern half of the ogdc. the low-lying area in the foreground and middle distance is underlain by the ogdc. the whitish parts of the low cliff across the fjord consist of augite syenite, whilst the brown-weathering rocks farther to the right are syenogabbros of the southern border group. the ridge in the middle distance behind the low cliffs consists of julianehåb batholith rocks. the high cliff in the far left distance consists of syenites of the tugtutôq central complex. 1818 that crystallised from below upwards with whole-rock compositions approximating to those of melts. as will be described below, there is evidence that compositionally stratified chambers played an important role elsewhere in the gardar province. it is a matter of speculation as to whether continuation of the differentiation trend in the hidden uppermost facies of the intrusion (below narsaq sund) led to more extreme agpaitic fractionates. a schematic vertical section of the ogdc is presented in fig. 9. thus, from the start of the evolution of the tuttutooq–ilimmaasaq–narsarsuaq lineament, the ogdc provides evidence bearing on the probable evolution of the peralkaline ilímaussaq suite. fe2++mnmg ca mn mn fe2+mg rims of pyroxenes in augite syenites and pulaskites pyroxenes olivines cores of pyroxenes in augite syenites and pulaskites fig. 7. pyroxene and olivine compositions in the ogdc plotted in terms of mg–(fe2+ +mn)–ca and mg–fe2+–mn respectively. modified from upton et al. (1985). fig. 8. transition between the syenogabbroic marginal zone (brown) and augite syenite of the central zone (white) in the ogdc. the crudely vertical elongation of the syenite facies may denote channelways within the thermal boundary layer through which low-density, residual trachytic melts ascended. fig. 9. schematic vertical section of the ogdc at tuttutooq (upton et al. 1990). the marginal gabbroic rocks are indicated in close stipple ornament whilst the syenites of the central zone are shown in light stipple. the sill-like expansion of the ogdc at the unconformity in the host rocks is hypothetical, as is the synformal layering in the deeper parts of the central zone. the wavy horizontal lines ‘a-a’ and ‘b-b’ diagrammatically indicate the different erosion levels at the shallowest (easternmost) and deepest (westernmost) outcrops, respectively. the width of the giant dyke is 0.5 km. the height difference between ‘a-a’ and ‘b-b’ is estimated at 2–3 km. a a b b julianehåb batholith eriksfjord formation 19 the younger giant dyke complex this massive gabbroic intrusion, intruded at c. 1163 ma (table 1), has the form of a bifurcating suite of giant dykes (upton 1962, 1964a, 1987; upton & thomas 1980). it crosscuts the ogdc and the time lapse separating these two intrusions probably amounted to around 20 million years. a change in trend between the two suggests a slight anticlockwise reorientation of the regional stress field. palaeomagnetic data show an apparent polar movement of c. 10° to the east between the two so that significant plate movement may have intervened (piper 1976). the younger giant dyke complex (ygdc) can be followed for c. 140 km from the labrador sea to the inland ice (figs 10–12). above the unconformity separating the batholith and the eriksfjord formation the dyke morphology switched to lopolithic much in the same manner as the muskox intrusion in arctic canada (irvine & baragar 1972). the branching pattern of the ygdc on the tuttutooq archipelago west of ilimmaasaq and the generalised attenuation of its branches from east-northeast to west-south-west suggest lateral flow of magma from a focus to the east-north-east. that this focus lay in the region of illimmaasaq is also indicated by the gravity map (fig. 4). however, both the giant dykes and the succeeding main swarm of the more fractionated postygdc dykes are traceable with undiminished intensity east-north-east of illimmaasaq to where they pass beneath the inland ice. at their maximum the dykes attain widths of 800 m although more generally they are 300 to 500 m broad. the entire intrusion of basaltic magma is deduced to have taken place during a single dramatic, large-scale, trans-tensional event. the mean initial magma composition, as indicated by analyses of what appear to be the least contaminated of the medium-grained doleritic marginal samples of the giant dyke branches close to the inland ice, closely matches that gained from study of the giant dykes on tuttutooq (upton & fitton 1985). this observation supports the hypothesis that intrusion of all of the ygdc occurred simultaneously and involved a very large and homogeneous magma batch. heat loss is assumed to have taken place principally through the walls. the magma crystallised as a closed system although late-stage generation of silica oversaturated salic magma in eastern tuttutooq probably involved crustal assimilation. although the dykes are dominantly composed of troctolite, syenogabbros, ferrosyenites and syenites (both silica oversaturated and undersaturated) occur in localised differentiated facies. plots showing compositional data on the ygdc olivines, feldspars and pyroxenes are presented in fig. 13. although the cooling of the ygdc magma was sufficiently slow for most of it to crystallise as coarse-grained troctolite, it was sufficiently rapid to inhibit migration of intercumulus melts, thus preventing the textural and/or chemical re-equilibration of the high-temperature prod 1 2 3 major elements (wt.%) sio2 43.47 46.00 46.06 al2o3 15.65 16.71 17.13 fe2o3 † 15.57 14.91 14.21 mgo 4.76 5.93 6.12 cao 7.70 7.78 7.96 na2o 3.45 3.55 3.43 k2o 1.81 1.45 1.41 tio2 4.40 2.63 2.51 mno 0.19 0.19 0.18 p2o5 1.95 0.86 0.83 total 98.95 100.01 99.84 trace elements (ppm) ni 22 52 63 cr 5 30 42 v 138 160 168 sc 16 17 17 cu 40 39 44 zn 81 91 90 sr 1039 901 921 rb 39 23 25 zr 162 150 141 nb 31 22 21 ba 1669 1120 1052 la 47 44 27 ce 103 64 63 nd 57 34 32 y 35 27 26 feo*/(feo* + mgo) wt.% 0.75 0.69 0.68 al2o3/cao 2.03 12.15 2.15 k/rb 385 522 468 ba/sr 1.651 1.24 1.14 zr/nb 5.2 6.8 6.7 la/y 1.34 1.63 1.04 1. older giant dyke, tugtutôq (chilled facies). n = 3.§ 2. younger giant dyke, tugtutôq (chilled facies). n = 9. 3. giant dykes, nunatak region and north-east of motzfeldt. n = 9. table 2. giant dyke compositions †total iron as fe2o3 or feo. §n = number of analyses. ygdc, gabbroic facies older gardar rocks and julianehåb batholith post -ygdc intrusions 10 km ygdc, differentiated facies geologfjeld illu tali k itillip saqqaa itillip saqqaa kange rlua nars aq sun d syenitknold tretungegletscher mellemlandet johan dahl land nordtop g.f. holm nunatak n sejlfjord itillinnguujukitillinnguujuk tullerunnat killiit sydtungegletscher m otzfeldt sø tuttu tooq brede fjord narsarsuaq qassiarsuk narsaq qooqqup sermia qo or oq sko vfjo rd tun ulli arf ik fig.3 upton/garde 29-08-2013 46° 61° 2020 ucts that adds complexity to many ‘classic’ layered intrusions. consequently, the troctolites and their associates are typical orthocumulates in which melt trapped within the interstices underwent extended in situ crystallisation. this yielded discrete intercumulus minerals and added zonal increments to the cumulus until the solidus was reached (wager et al. 1960). some subsolidus migration of aqueous fluids is also inferred. fig. 10. distribution of the younger giant dykes in the east-northeast of the gardar system, extending (left to right) from kangerlua to g.f. holm nunatak, and from the outer islands of the tuttutooq archipelago to narsaq sund and their higher-level representatives around narssaq and the ilímaussaq complex. narrow extensions of the ygdc continue 15 km south-west of the map boundary. composition of the magma the mean compositions of the chilled marginal facies of the ygdc on tugtutoq and in the more easterly giant dyke branches, and the chilled marginal composition of the ogdc, are shown in table 2. the initial magmas of the ygdc and its ogdc predecessor are inferred to have been closely related. both had compositions close to ‘the critical plane of undersaturation’ (yoder & tilley 1962) but the ygdc magma had higher alkalis and marygdc, gabbroic facies older gardar rocks and julianehåb batholith post -ygdc intrusions 10 km ygdc, differentiated facies geologfjeld illu tali k itillip saqqaa itillip saqqaa kange rlua nars aq sun d syenitknold tretungegletscher mellemlandet johan dahl land nordtop g.f. holm nunatak n sejlfjord itillinnguujukitillinnguujuk tullerunnat killiit sydtungegletscher m otzfeldt sø tuttu tooq brede fjord narsarsuaq qassiarsuk narsaq qooqqup sermia qo or oq sko vfjo rd tun ulli arf ik fig.3 upton/garde 29-08-2013 46° 61° 21 despite these overall undersaturated characteristics, silica-oversaturated rocks occur in eastern tugtutôq in the vicinity of asorutit (see below). these may represent a local anomaly due to crustal assimilation. a photomicrograph of a chilled ygdc marginal sample is presented in fig. 14. ndand sr-isotopic features are close to bulk earth values (upton et al. 2003). 87sr/86sr1163 values for the ygdc troctolites range from 0.70279–0.70321, suggesting insignificant crustal contamination (mingard 1990). δ18osmow values for feldspars from the troctolites are +5.0 to +6.5‰; troctolitic and peridotitic whole-rocks give δ18osmow values of +4.0 to +5.4‰ and +4.4 to +6.9‰, respectively. feldspars from the assorutit quartz syenite have δ18osmow of +6.5 to +6.8‰ and the values for the corresponding whole-rocks range from +4.2 to +5.9‰. the oxygen isotope values for the troctolites are slightly lower than those of fresh skaergaard gabbros (taylor & forrester 1979) where the feldspars have δ18osmow +7.6 to +8.7‰ and whole-rocks +7.2 to +8.2‰. mingard (1990) considered that the low ygdc values (<+5‰) could be due to localised interaction between magma and meteoric water but noted that assimilation of low δ18o lower crust could not be excluded. crystallisation sequence melting experiments at 1 kb suggest that the magma was intruded at 1140° ± 10°c, in equilibrium with olivine and plagioclase, whilst the solidus was at c. 980°c (upton 1971). petrographic evidence indicates delayed crystallisation of augite and experimental (1 kb) work indicates that the olivine–plagioclase–clinopyroxene–liquid cotectic was not attained until 1060°c ± 15°c. the layered cumulates in the ygdc indicate that crystallisation of fe-ti-oxides and apatite preceded clinopyroxene. studies on associated fine-grained dykes (the main swarm described below) suggest that fe-ti-oxides and apatite joined the assemblage when mgo in the liquid had fallen to c. 3.75 wt% whereas pyroxene phenocrysts did not appear until this value had been reduced to c. 3.25 wt% (upton & thomas 1980; martin 1985). such relatively delayed crystallisation of pyroxene is unusual in basaltic systems. crystallisation of fe-ti-oxides and apatite appears to have taken place within a very narrow temperature interval. the ygdc magma was relatively reduced with an oxygen fugacity lying between the qfm and iw buffers (upton & thomas 1980). ginally lower k2o/na2o (0.41 as opposed to 0.48). its notably high k2o content (1.43 wt %) is regarded as a primary characteristic, unrelated to crustal assimilation. the ygdc magma had lower cao, al2o3, tio2 and p2o5 contents than the initial ogdc magma and was also poorer in incompatible trace elements, especially ba and sr. the f content was approximately half that of the ogdc magma (upton & thomas 1980; köhler et al. 2009). the ygdc magma was relatively poor in normative diopside and consequently crystallised to troctolitic rocks (upton & thomas 1980; upton 1996). it was marginally silica undersaturated (c. 2% normative nepheline); residual veinlets of nepheline syenite are known from several localities at tuttutooq, and a substantial body of nepheline syenite is present within the most easterly ygdc extensions in the nunatak region (described below). 2222 fig. 12. oblique aerial photograph along the ygdc, looking east-north-east along tuttutooq. the northern part of the ilímaussaq complex (with ice) is seen in the far distance. the valley to the right is excavated from ygdc gabbro. the parallel grooves in the foreground denote weathering of main swarm dykes and shear zones in the julianehåb batholith. pale-coloured outcrops on the ridge in the middle distance are due to fast-weathering, crumbly outcrops of the ygdc. fig. 11. view towards east-north-east along the northern branch of the ygdc in western tuttutooq, showing typical hummocky topography. the high ground to either side and in the distance is underlain by the julianehåb batholith. 23 fig. 13 (modified from mingard 1990). a: olivine compositions in the younger giant dyke complex, shown in the fo–fa–tp ternary diagram, from fo68 to near-end-member fayalite but with a small late-stage increase in tp. b: feldspar compositions in the ab–an– or ternary, extending from an65 through more sodic plagioclases and ternary compositions to alkali feldspar close to the minimum melting composition on the ab–or join. c: clinopyroxenes in the system en–fs–wo, grading from salite to ferrosalite. in syenitic differentiates more extreme compositions (not shown) extend to ferrohedenbergite and aegirine-augite. a b c forsterite albite fayalite orthoclase wollastonite anorthite tephroite enstatite ferrosilite internal structures although the greater part of the intrusion consists of homogeneous troctolite, layered cumulates and/or differentiated rocks appear at irregular intervals along the dyke branches. layering features define synformal structures which dip symmetrically inwards from the sides to horizontality along the dyke axes. where closure can be discerned the layering has a canoe-like morphology, dipping inwards from the ends as well as from the sides (fig. 15). the layered parts of the dyke may be described as nodes or pods inferred to mark sites, commonly only a few hundred metres long but which can be up to 3 km, where convection cells developed. the observation that the layering varies from barely visible to strongly accentuated may relate to the vigour of convection. the relatively abrupt changes along the intrusion from isotropic troctolite to the dozen or more layered nodes imply that the factors dictating the change from the one to the other were critically poised. these factors are presumed to have included the rate of heat loss, the local morphology of the intrusion, melt composition, and possibly structural depth in the intrusion (irvine 1987). the presence or absence of cumulate layering was not simply governed by dyke width as there are broad sectors lacking cumulate features and narrow sectors in which layering is strongly developed. an astonishing variety of layering styles is exhibited, including feldspar lamination alone, normally-graded rhythmic layering, isomodal layering, micro-rhythmic layering, diffuse modal layering and graded rhythmic layers alternating with uniform ‘standard’ rock. the ygdc in the tuttutooq archipelago in view of the spectrum of phenomena relating to fluid dynamics in the convecting cells in the magma chamber, the key localities along the tugtutôq dyke branches are here described below from east-north-east to westsouth-west. the descriptions are mainly based on upton (1964b, 1987), upton & thomas (1980), and upton et al. (1996), and some observations have not been published before. figure 16 shows a series of schematic crosssections along the ygdc. 2424 sissarluttooq the southern branch of the ygdc reaches its greatest observable thickness (c. 800 m) at the extreme east-southeast corner of tuttuooq, on the sissarluttooq peninsula (fig. 17). however, the divergence of contacts towards the ene implies its further widening beneath the waters of narsaq sund. a coastal section on its northern side reveals the chilled marginal zone transecting the ogdc whereas the southern margin is unexposed. the principal feature of interest at sissarluttooq lies in the well-developed lamination due to parallel orientation of idiomorphic plagioclase crystals 2–3 mm across, tabular parallel to the (010) faces. olivine is the secondmost abundant component (up to 1 mm diameter) whilst idiomorphic apatite and titanomagnetite crystals are also cumulus components. although modal layering is abfig. 14. photomicrograph of the chilled margin of the ygdc against julianehåb granite. the opaque, flow-banded, chilled zone inferentially reflects a vitric facies subsequently recrystallised and oxidised. pale spots in the lower part of the image are presumed spherulites. the original texture of the granodioritic country-rock has been profoundly modified by partial melting and recrystallisation. two-way element exchange is assumed, involving entry of k and other mobile elements from the country-rock into the hot, chilled gabbro. back-veining from the host rock has not occurred. horizontal field of view 30 mm. 15. schematic sketch of a layered synform in the ygdc. increasing thickness of ferromagnesian cumulate towards the centre is indicated diagrammatically. 1 2 3 4 5 6 500 m 7 troctolitic gabbro modal and/or lamination layering julianehåb batholith fig. 16. cross-sections through the younger giant dyke at seven localities. 1: itillinguujuk. 2: tripyramidal peak west-south-west of itillip saqqaa. 3: itillip saqqaa. 4: marraat. 5: sissarluttooq. 6: krydssø. 7: syenitknold. locations: 1–6 on tuttutooq and its western islands; 7 on nunatak north of motzfeldt sø, see fig. 10. 25 sent, the lamination in the central part of the intrusion (50–100 m from the contacts) defines synformal layering with inward dips of c. 40°, decreasing axially. the structure, however, is not that of a simple single basin as there appear to be several foci at which the lamination attains horizontality. the total length of the laminated pod is c. 1 km. west-south-west of sissarluttooq, the southern branch reverts to homogeneous troctolitic gabbro, but at a distance of c. 4 km it subdivides into a layered northern sub-branch and a homogeneous southern sub-branch. in the former, layering is shown by feldspar lamination that dips inwards at c. 30° before shallowing symmetrically to zero along the axis. this synformal structure persists for about 1 km before the gabbro dyke resumes homogeneity. marraat two kilometres west-south-west sissarluttooq in the southern sub-branch, the coastal section at marraat provides excellent exposures across the 500 m wide dyke. for the first few metres in from the contacts, thin vertical mafic layers occur parallel to the contacts with planar alignment of plagioclases. it is within these border-group rocks that the phenomenon of ‘snowflake’ plagioclase glomerocrysts occurs (fig. 18). in the axial 300 m of the dyke, modal layering with feldspar lamination becomes more strongly developed with inward dipping layers at increasingly low angles to the central horizontal zone. mafic layers up to 10 cm thick are due to concentrations of olivine, titanomagnetite and apatite; indications of normal grading and some cross-bedding are indicative of magma flow. asorutit the northern branch of the ygdc crops out between the nasaasarli and asorutit headlands (fig. 17). like its southern counterpart, this branch is c. 800 m broad but the contact zones are now parallel, rather than divergent as in the former. apparently homogeneous gabbros form 200 m thick marginal border groups that grade, through a few metres of iron-rich syenogabbro, into a vertical zone several metres broad, displaying a spectacular array of directionally oriented, branching, clinopyroxene n 250 140 240 230 110 170 115 170 arfersuaq asorutit nasaasarli 305 krydssø sigssarluttoq troctolitic gabbro krydssø syenogabbro assorutit quartz syenite younger gardar dyke complex older gardar dyke complex foyaites julianehåb batholith section a–a’ section b–b’ kryd ssø 1 km b b’ a a’ a b 46°15´ 61° 60°54´ fig. 17 (modified from upton 1962, 1964a). a: geological map of the south-eastern coast of tuttutooq, showing two ygdc branches cutting across the ogdc. the narrowing of the southern branch between the 230 and 240 m spot-heights reflects upward narrowing of the intrusion and is topographically controlled; this is not the case for the narrowing of the northern branch between krydssø and asorutit. the central, differentiated facies in the asorutit area is composed of quartz syenite, separated from the marginal troctolitic gabbro by narrow syenogabbroic margins. the central, differentiated rocks at krydssø comprise synformally layered syenogabbros grading up into augite syenite. little or no layering is visible. b: enlarged, schematic cross-sections of the differentiated facies at asorutit and krydssø. the synformal layering at asorutit is hypothetical. fig. 18. glomeroporphyritic clusters of plagioclase and olivine at marraat, composing ‘snowflake’ textures. pocket knife 10 cm long. 2626 dendrites up to 50 cm long (fig. 19). it is deduced that inward crystallisation of the giant dyke enclosed residual magma in which the pyroxene components became increasingly concentrated until a critical degree of supersaturation was reached, when nucleation of the dendritic crystals was triggered to form these remarkable crescumulate zones, symmetrically developed on either side of the intrusion. sandwiched between these transitional syenogabbro border zones is a central body of syenite. in plan, this wedges out sharply as traced westwards, terminating shortly before the dyke narrows to a ‘wasp waist’ less than 250 m broad (fig. 17). the eastern outcrop of the syenite disappears beneath sea-level but, by analogy with salic cores elsewhere in the giant dykes both in the tuttutooq–ilimmaasaq–narsarsuaq system and in the isortoq region, the syenite is probably a localised lenticular development less than 2 km long. there are no clear indications of any layering features other than localised ferromagnesian-rich schlieren. the texture, however, suggests that the syenite is an orthocumulate (fig. 20) and it is speculated that the visible rocks are underlain by layered cumulates comparable to those of krydssø (described below). the syenite is hypidiomorphic granular, composed c. 80% modally of squat rectilinear alkali feldspars 7–8 mm long. cryptoperthitic cores grade out to more coarsely exsolved antiperthites, surrounded by clear outer zones of albite. other early phases are ferroaugite, fayalitic olivine, titanomagnetite and apatite. colourless to grey clinopyroxene cores zone out to pale green and, locally, to deeper green aegirine-augite. the olivine (fo5–2.5) contains exsolved parallel plates of fe-oxide with variable replacement by iddingsite. reaction fringes of blue-green alkaline amphibole grading sharply out into arfvedsonite surround the pyroxenes and olivines. this discontinous reaction series culminates in biotite, zoned from strongly pleochroic to colourless, as the youngest ferromagnesian mineral. the latest (intercumulus) components are quartz and calcite. whereas the asorutit syenite and the more primitive syenites of the ogdc share many petrographic affinitities, the development of the albite coronae around the perthites and the intercumulus quartz and calcite distinguish it from its silica undersaturated ogdc predecessors. the youngest components of the ygdc suite at asorutit are localised patches of granite pegmatite in the syenite. veins of this coarse alkali granite (perthite, quartz, arfvedsonite and accessory zircon) are prominent over a distance of c. 2 km along the northern ygdc branch and are regarded as filter-pressed residues from the latter. krydssø on the western side of the northern branch constriction, the giant dyke again appears composite but now with a core of syenogabbro/ferrosyenite (the krydssø body) rather than a leucocratic syenite (fig. 17). the core is lenticular in plan, c. 3 km long and, like the assorutit syenite, it is ensheathed by troctolitic border groups. there is a degree of symmetry in plan between the krydssø and the asorutit bodies: in the central part of the giant dyke at krydssø the rocks pass upwards from syenogabbro with fig. 20. coarse-grained quartz syenite at asorutit. the interstices between idiomorphic perthitic alkali feldspar crystals mostly contain sodic amphibole, quartz and calcite. diameter of coin 2.5 cm. fig. 19. dendritic clinopyroxene prisms in the syenogabbroic zones separating the asorutit quartz syenite from the border-group gabbros. distance across the outcrop left to right is c. 75 cm. 27 anhedral, intercumulus, pyroxene to syenites containing idiomorphic prismatic augite. it is the only place in the ygdc where phase layering has been observed rather than inferred. whilst its crystallisation may have been coeval with that of the asorutit syenite, the krydssø syenogabbro/ferrosyenite body is transected by some of the residual alkali granite veins. contact relations between the iron-rich core and the sheath of enveloping gabbro are unexposed. modal layering in the syenogabbro dips inwards around the body so that any one layer thus defines a boat-like morphology. the modal layering involves alternation of melanocratic and mesocratic layers on a decametre scale, both with and without normal grading (fig. 21). the krydssø syenogabbros present one of the few opportunities along the ygdc to observe the axis of the synformal layered structure (fig. 22). modal layering appears in otherwise homogeneous troctolite along the southern coast of store pilesø as ‘inch-scale layering’ i.e. micro-rhythmic isomodal alternation of mafic and felsic layers, each only 2–3 cm thick, presenting a unique layering style in the intrusion. beyond this, as traced west-south-west, the northern branch is devoid of cumulate layering for nearly 16 km; this homogeneity persists until just west of the dislocation of the northern branch by one of the left-lateral faults that cut the southern rift. in the offset intrusion, fig. 21. normally graded modal layering in ferro-syenogabbro, krydssø. scale is 50 cm long. fig. 22. view towards east-north-east beside krydssø along the axial plane of the layered synform. ferro-syenogabbros in the foreground dip nnw, while those in the distance dip sse (white lines). 2828 south of the fault at itillip saqqaa, the most striking layering phenomena appear in a style that is unique within the gardar province. itillip saqqaa the ygdc northern branch has been displaced 1300 m sinistrally by the fault; the gabbros on the immediate southern side of the fault are below sea-level so that the nearest outcrops for study are from 250 to 800 m distant from the fault plane on the itillip saqqaa coast (fig. 23). although the transcurrent fault at itillip saqqaa has clearly displaced the giant dyke, marked contrasts in the dyke features on either side of the fault strongly suggest that some pre-existing structural element affected crystallisation prior to the main left-lateral movement. north-east of the fault plane the troctolite gabbro is homogeneous, contrasting with distinct heterogeneity imposed by modal layering on its south-western side. evidence for convection, high-temperature faulting and itillip saqqaa 270 n 0.5 km giant dyke complex julianehåb batholith older gardar dyke (bd0) syenitic differentiate troctolitic gabbro 46°30´ 60°50´ fig. 23. geological map of giant-dyke relationships at itillip saqqaa on central tuttutooq, showing members of the ogdc, ygdc and an older gardar dyke (bd0) in the south-west. the wnw–esetrending contact of an offshoot from the ygdc (marked by an arrow) is parallel to the fault at itillip saqqaa and is inferred to have been controlled by a pre-existing shear zone. fig. 24. intermittent spacing of thin black peridotitic layers in grey troctolites (itillip saqqaa). scale is 50 cm long. fig. 25. ‘snowflake’ glomerocrysts in a small (c. 2 m wide) dyke. small island off the south-east coast of tuttutooq. scale is 50 cm long. 29 gravity slumping, present on the south-west side of the fault, is lacking to the north-east and there are no indications of the dyke parting into two branches. the possibility of major vertical displacement along the fault cannot be discounted. if so, downthrow to the north might be inferred, with the giant dyke outcrops to the south representing deeper structural levels. it is suggested that a plane of weakness, pre-dating the sinistral displacement, was already extant. prominent layering in the troctolites south of the fault typically consists of: (a) whitish-grey mesocratic troctolites with c. 70% (modal) plagioclase and c. 30% olivine (less than 10% interstitial augite and titanomagnetite), and (b) nearly black feldspathic peridotite consisting of >80% olivine (fig. 24). the mesocratic rocks are considered to represent products of crystallisation along an olivine-plagioclase-liquid cotectic, with c. 30% olivine and c. 70% plagioclase. typically there are sharp contacts between the contrasting pale and black layers. in some places the peridotite layers display irregular bases, attributed to differential loading by dense layers of olivine cumulus overlying readily deformable feldspar-rich layers. the black colouration of the peridotites is due to microscopic opaque inclusions in the olivines. the magma is inferred to have reached a shallow crustal level whilst supersaturated with respect to plagioclase. reduction in pressure on ascent is inferred to have stimulated plagioclase nucleation. the plagioclase grew rapidly, commonly from olivine nuclei, to form radiating ‘snowflake’ glomerocrysts (berg 1980). the resultant troctolites lack plagioclase lamination and are inferred to be wholly composed of polymineralic ‘snowflakes’ that accumulated as cumulus. it is, however, only in chilled marginal facies that the ‘snowflake’ morphology becomes apparent. well-developed large-scale ‘snowflake’ clusters are shown in a related dyke rock, outside the ygdc (fig. 25). the occurrence of dendritic plagioclase growth normal to peridotite layers can be seen in this part of the ygdc (fig. 26). as with the growth of ‘snowflake’ clusters, it provides evidence that episodic supersaturation of the magma with respect to plagioclase was relieved by rapid growth of feldspar from a boundary layer to form a perpendicular feldspar crescumulate (wager et al. 1960). accumulation of the ‘snowflake’ glomerocrysts produced the troctolitic layers, the fabric of which consequently differs from that of the well-laminated troctolites (e.g. as seen at sissarluttooq). intermittently plagioclase failed to nucleate, leaving olivine crystallising alone to form the peridotites. thus the bimodal layering resulted from whether or not plagioclase was crystallising. the peridotitic layers are also characterised by a characteristic jointing that is either normal to the layering or at a high angle to it. two parallel synforms are present in the ygdc cross section at itillip saqqaa. in the narrower and steeper southern synform, cross-lamination as well as some degree of normal grading between troctolite and peridotite is present. these features are attributed to erosion and deposition, together with crystal winnowing, resulting from vigorous downflow of magma adjacent to the southern contact of the giant dyke (fig. 27). more fig. 26. white dendritic plagioclase grown upward from and within peridotite layers in the ygdc west-south-west of itillip saqqaa. sections on scale are 5 cm long. 3030 tranquil deposition of olivine and olivine + plagioclase clusters appears to have typified the broader northern synform. peridotite layers tend to thicken down-dip. older (lower) layers dip more steeply than stratigraphically younger layers in the southern synform. this structure exhibits a miniature equivalent to oceanic dipping reflectors in which the oldest units have the steepest dip towards the volcanic zone whilst younger units have progressively lower dips; the processes occurring in the two scenarios are thought to have been comparable: as crustal extension persisted in the dilating dyke fissure, the earlier layers rotated downwards to be progressively overlain unconformably by younger layers. normal faults developed at high temperatures (above 600°c?) when the cumulates acquired a capacity for brittle fracture. the faults throw down towards the synform axis so that the central zone was undergoing ‘graben’ subsidence (fig. 28). whereas opening of the giant dyke fissures took place sufficiently fast to form deep, steep-walled magma chambers, this feature demonstrates that dilation of the dyke was still taking place while cumulates were being depositfig. 27. normally-graded and cross-bedded layers in the southern part of the paired synform at itillip saqqaa. the dark unit (c. 2 m thick) behind the person in the right is massive peridotite. fig. 28. five peridotite layers in troctolite that have undergone ductile deformation in a normal fault with downthrow towards the layered synform axis. itillip saqqaa. scale is 50 cm long. 31 ed. thus the extensional stresses were still being exerted, not only during the deposition of the cumulates but also during the subsequent cooling history of the intrusion. in the northern synform the thickness ratio of troctolitic versus peridotitic layers decreases from flanks to centre, i.e. the marginal parts of the synform are predominantly mesocratic (fig. 29) whilst those of the axial region are dominantly melanocratic (fig. 30). comparable downdip thickening of mafic/ultramafic layers occurs in other gardar cumulate bodies, e.g. within the kûngnât fjeld syenites (upton et al. 2013). there is a notably inequigranular (i.e. non-equilibrated) texture shown in the peridotites (fig. 31). near the northern synform axis a slump breccia occurs, comprising clasts of peridotite up to half a metre across, enveloped in a troctolitic matrix (fig. 32). some of the clasts show layering, and the angular discordances from clast to clast clearly indicate their rotation during slumping. deformed layering in the surrounding matrix points to it having been highly ductile or mushy, contrasting with the coherent clasts. the form of many of the peridotitic clasts is defined by roughly planar surfacfig. 29. ygdc outcrop close to itillip saqqaa showing thin peridotite layers within troctolite towards the margin of northern synform. height of outcrop c. 3 m. fig. 30. synform axis, itillip saqqaa. in contrast to fig. 29, this outcrop displays thick peridotite layers, separated by thinner troctolite layers more resistant to weathering. scale bar is 0.5 m long. 3232 es, suggesting that a joint system had already developed at the time of their disruption, and it is presumed that this corresponds to the jointing seen in the peridotite layers. hence the peridotites had already achieved a state capable of brittle fracture while the troctolites were still readily deformable. the situation mirrors that shown by ultramafic and feldspathic material in syenite cumulates at the nunarssuit and kûngnât complexes (upton et al. 1996). the slump breccia at itillip saqqaa is deduced to have formed when steeply dipping (jointed) peridotite layers in the synform limbs became gravitationally unstable and collapsed, yielding a chaotic breccia in the synform hinge-zone. remarkable features, specific to the giant dyke between the coast at itillip saqqaa and the tripyramidal peak 3 km farther west-south-west, were given the fieldname the ‘troll’s smile’. these features comprise crescentshaped peridotite bodies between 5 and 10 m long and up to 3 m thick, concave-up (as seen in dyke-parallel exposures) within the pale-coloured troctolites. the crescents exhibit a crude columnar jointing, normal to their margins, which gives the fanciful appearance of a smiling mouth whilst the jointing gives the impression of irregular teeth (fig. 33). these features are interpreted to be cross-sections of fan-like olivine-rich bodies that propagated downwards and inwards from the dyke sidewalls at c. 90° to their strike. to explain the ‘troll’s smile’ phenomenon it is proposed that snowflake cumulus cascaded continuously down the thermal boundary layers to accumulate on pre-existing crystallising cumulus, to form a ‘sedimentary’ pediment sloping down to the median axis. this process would have been unsuspected had it not been for the sidewall foci where plagioclase nucleation failed, leaving olivine to crystallise and sink alone. steep localised channels leading down normal to the dyke walls can be inferred, down which dense slurries of melt + olivine crystals would have flowed. the slurries initially excavated troughs in unconsolided cumulus bebfig. 31. photomicrograph (plane polarised light) of typical peridotite from itillip saqqaa. note the un-equilibrated texture and the large inequality of sizes of the olivine cumulus. intercumulus minerals are plagioclase, magnetite and scarce augite. olivine crystals are up to 2 mm across. fig. 32. slump breccia of peridotite clasts in troctolite matrix. itillip saqqaa. hammer length c. 26 cm. fig. 33. a: lenticular-section of a peridotite unit in the western ygdc. the peridotite is abruptly defined between troctolite above and below, and displays characteristic jointing transverse to cooling surfaces. scale-bar (centre, against peridotite) is 50 cm long. b: sketch showing joints in the peridotite (‘troll’s smile’). 33 fore depositing their load of olivine crystals as velocities decreased. this formed peridotitic deposits that widened outwards toward the dyke axis. morphologically such features may be compared to sedimentary alluvial fans (fig. 34). although no size analyses of the olivines have been made, the mean size of the crystals may be expected to increase downslope towards the axis. tripyramidal peak west-south-west of itillip saqqaa three kilometres west-south-west of itillip saqqaa, the northerly ygdc sub-branch underwent localised expansion (‘ballooning’) before abruptly narrowing to a third of its thickness (fig. 10). the expanded section is less than 1 km long and gives rise to a characteristic tripyramidal peak. the reason for the thickness change is unknown; whereas the expanded section is composed of homogeneous troctolite, the thinned section west of it is strongly layered, exhibiting the same style of layering as at itillip saqqaa. in this section the limbs of a layered synform contains jointed peridotite layers that thicken down-dip (fig. 35). this phenomenon is attributed to gravitational creep or saltation (‘jumping’) of cumulus olivines. itillinnuujuk the northern sub-branch of the northern ygdc branch has a wedge-shaped termination a few kilometres farther west but then reappears on the island of itillinnguujuk, where it is reduced to a width of 200 m (fig. 10). the locality is of interest in showing a repetition, albeit in a narrower dyke section, of the features seen at marraat. as at the latter, synformally layered gabbro (defined by modal layering) forms the central third of the dyke with the enclosing border groups composing the other two thirds. the radial growth (‘snowflake’ texture) of plagioclase fig. 34. schematic section of the giant dyke illustrating the concept of localised ‘sedimentary fans’ descending from the side walls to spread out as they approach the synform axis. fig. 35. peridotite layer with oblique jointing, exhibiting characteristic down-dip thickening towards the synform axis. 3434 around olivine nuclei is again well shown. the locality is the most extreme westerly one at which a synformally layered ‘pod’ is developed within the ygdc. figure 16 shows the variability in the cross-sections of the ygdc along its outcrop. minor offshoots from the giant dykes there are many smaller basaltic dykes parallel to the ygdc that are regarded as offshoots from it (fig. 36). a notable example is seen where the magma of the southern sub-branch of the southern ygdc branch, inferred to have been propagating in a westerly direction, encountered a crush zone in the julianehåb granite. here it was not completely stopped but continued westwards beyond the crush as a comb-like swarm of a dozen or more small (<5 m) dolerite dykes that can be traced for several tens of kilometres. macdonald et al. (2010) presented an argument for some of the large palaeogene dykes in southern scotland having been arrested at fault planes that were acting as aquifers. following the same line of reasoning it is suggested that, assuming the giant dyke branch was being propagated with a strong west-south-westerly component, the halting of the bulk of the magma resulted from water-cooling at the crush zone. a portion of the magma, however, was able to penetrate beyond it, forming the dyke swarm. narsaq gabbro and lopolithic relicts the mildly alkaline gabbros cropping out at narsaq and along the nuugaarmiut peninsula to its north were referred to as essexites by ussing (1912) and as essexite gabbros by wegmann (1938). the outcrop has a generalised nw–se trend for some 3 km from the tip of the nuugaarmiut peninsula to fabriksbugt (fig. 37). the petrography and internal structures have such close affinity with those of the tugtutôq giant dykes (c. 5 km to the wsw) as to remove doubt that these gabbros are integral components of the younger giant dyke complex. to the west, the gabbro is bounded by the waters of the narsaq sund while to the east, it is truncated by the narssaq syenite (fig. 37). critical information on the relationship of the intrusion to older formations comes from two exposures on both sides of niaqornaarsuk. the vertical contact between gabbro and julianehåb granite (fig. 38) does not have the regional ene–wsw trend but is oriented be1 km asorutit sissarluttoq narsaq sund nuugaarmiut narsaq niaqornaarsuk fab riks bug t tuttu tooq assorutit quartz syenite krydssø syenogabbro troctolitic gabbro strike and dip of layering diatreme mela-aillikite narssaq syenite narssaq gabbro julianehåb batholith younger giant dykes n 46°07´ 60°54´ fig. 36. typical small homogeneous dyke of dolerite. nasaasarli, east coast, tuttutooq. fig. 37. geological sketch map showing relationship of the giant dykes on eastern tuttutooq to the gabbro outcrop around narsaq and nuugaarmiut. 35 tween ese–wnw and e–w. it presents a well-chilled, fine-grained marginal facies with small subhorizontal pegmatitic segregations persisting for several metres from the contact. this contact is correlated with the southern contact of the ygdc on sissarluttooq, described above. at fabriksbugt a contrasting exposure reveals a chilled contact of the gabbro dipping at a low angle against quartzite strata of the eriksfjord formation (fig. 39). the fine-grained gabbro is crowded with plagioclase megacrysts and small xenoliths of anorthosite. over the unexposed 1200 m between these two exposures the geometry of the ygdc clearly changed from that of a vertical dyke in the julianehåb granite to a subhorizontal sill or lopolith at the unconformity between the granite and the eriksfjord formation. in the vicinity of narsaq, away from the contact zones described above, the gabbro is coarse (c. 1 cm) and mainly structureless, lacking lamination or modal layering. anorthosite xenoliths, mostly <1 m, occur plentifully together with plagioclase megacrysts several centimetres across (fig. 40). many of the critical outcrops on glaciated slabs have later been covered by buildings as the town expanded. the gabbro in and around the oldest part of narsaq has been pervasively affected by hydrothermal alteration that caused sericitisation of the feldspars and replacement of the olivines, pyroxenes and biotites by chlorite and epidote. the alteration is ascribed to low-temperature fluids exuded during crystallisation of the cross-cutting fig. 38. vertical contact between gabbro (left) and granite (right of shadow in centre of image). west coast of niaqornaarsuk near narsaq. for location see fig. 37. fig. 39. chilled margin of the narssaq gabbro, crowded with plagioclase megacrysts and scarcer anorthositic xenoliths up to 5 cm large. fabriksbugt, narsaq. fig. 40. anorthosite xenolith (c. 25 cm across) and plagioclase megacrysts, narsaq township. 3636 narssaq syenite (described below). traced north-westward into the nuugaarmiut peninsula the alteration, and also the anorthositic xenoliths and associated megacrysts, disappear and the gabbros acquire a layered structure shown by plagioclase lamination and conformable, but discontinuous, mafic layers. the layers consist of concentrations of olivine, apatite and magnetite: in one such layer the mode was 12.5% plagioclase, 50% olivine, 30% titanomagnetite and 7.5% apatite. plagioclase within these mafic layers shows a higher degree of lamination than those in the average gabbro suggesting that density winnowing by flowing magma was responsible both for the concentration of heavy minerals and the lamination (upton 1961). sharp changes in dip and strike of the layering features and the truncation of the mafic layers by shear zones reflect penecontemporaneous tectonic or gravitational instabilities within incompletely solidified cumulates. the layering strikes approximately e–w with dips to the north of between 10° and 40° (fig. 37). assuming a generalised dip of 25°, the total stratigraphic thickness of these cumulates is roughly 530 m. whilst the more southerly, unlayered gabbros, with their content of anorthosite xenoliths and plagioclase megacrysts, are inferred to lie close to the intrusion roof, it is only at deeper levels, i.e. in the more northerly outcrops, that modal layering developed. between narsaq and the ilímaussaq intrusion there are several gabbroic masses which are inferred to be xenolithic relicts of former easterly extensions of the narssaq gabbro. the largest of these are on the 681 m mountain (qaqqarsuaq), to the east of the town and on the talut ridge close the western contact of the ilímaussaq intrusion. these are interpreted as part of one or more sheets, at least 100 m thick, the tops of which have been eroded (bridgwater & harry 1968 and references therein). fastchilled facies contain skeletal plagioclase phenocrysts, some as ‘snowflake’ aggregates (fig. 41). anorthosite, gabbro-anorthosite xenoliths and plagioclase megacrysts also occur as inclusions in the intrusions younger than the ygdc. some 8 km north-east of narsaq gabbroic anorthosite containing abundant feldspar megacrysts occurs as rafts in the ilímaussaq lujavrite on kvanefjeld, and isolated fragments of anorthosite and laminated gabbro are known from the ilímaussaq augite fig. 41. polished slab 6.5 cm wide of rapidly cooled dolerite at qaqqarsuaq with quenched, skeletal and concentrically zoned plagioclase megacrysts and phenocrysts. some of the latter occur as ‘snowflake’ glomerocrysts. accumulated (floated) anorthositic debris synformally layered cumulates, grading up into residuals julianehåb batholith eriksfjord formation fig. 42. hypothetical cross section of the giant dyke and overlying intrusions. not to scale. 37 syenite (hamilton 1964; sørensen 2006). the quantities of anorthosite xenoliths and plagioclase megacrysts in these occurrences support other observations that they came from high structural levels in the gabbro and had been concentrated as a result of flotation. whereas the great bulk of the original gabbroic intrusion, of which they are considered to have been a part, was engulfed by younger, syenitic and granitic magmas, these relatively low-density roofing facies containing feldspathic rocks failed to sink. figure 42 presents a hypothetical cross section, showing a giant dyke expanding into a lopolith at the julianehåb granite–eriksfjord formation unconformity, with overlying ‘cedar-tree’ laccolithic extensions into the overlying supra-crustal strata. whilst the original extent of this lopolith is a matter for speculation, the relics in the narssaq and ilímaussaq intrusions suggest that it may have measured some 10 km north-west–south-east and extended a similar distance north-eastwards. younger giant dyke extensions west and north of motzfeldt sø gabbroic giant dykes crop out from 25 to 60 km eastnorth-east of qassiarsuk in mellemlandet and the nunataks west and north of the motzfeldt complex (fig. 10). the widths, courses, compositions (table 2) and internal structures of these so closely resemble those on tuttutooq as to leave no doubt that they are easterly components of the same intrusion. these giant dyke extensions are, however, sinistrally offset (c. 20 km) from the tugtutôq dykes by faulting. two approximately e–w-trending, left-lateral faults, or fault systems, were responsible. the southernmost of these faults across the ilímaussaq peninsula controlled the emplacement of the narssaq complex described in a later section. displacement along this fault is inferred to have shifted the ygdc intrusion some 10 km to the west so that its outcrop is now almost wholly concealed by the waters of the bredefjord. this section of the ygdc would remain wholly hypothetical if it was not for a short section of gabbroic dyke cropping out a few kilometres west of qassiarssuk at kangerlua (fig. 10). another c. 10 km of translation to the west was probably brought about by a more northerly fault system that traverses the qassiarsuk and south qooroq areas (fig. 10). on the eastern side of northern tunulliarfik there is a deep gorge, co-linear with the giant dykes to the east-north-east in johan dahl land (walton 1965), the outflow from which contains troctolitic boulders. consequently the continuation of one of the giant dyke branches is inferred in this sector. three giant dyke branches traverse g.f. holm nunatak but only the southern branch of these is continuous south of nordtop from sydtungegletscher to syenitknold. from its absence on mellemlandet (to the wsw), this branch is presumed to pinch out beneath the qooqqup sermia glacier. the middle branch on g.f. holm nunatak is absent on the nordtop nunatak, having terminated to the west-south-west of g.f. holm nunatak. the northern branch is presumed to underlie the ice just north of nordtop, traversing the northern part of mellemlandet (fig. 10) to the wsw and on through much of johan dahl land, but failing to reach the fjord section north of narsarssuaq. the east-north-east extensions of the ygdc have vertical, typically parallel sides and maintain more-orless constant widths of 300 to 600 m. they can, however, attenuate to zero thickness over a distance of one or two kilometres as exemplified by the southern dyke branch on mellemlandet (fig. 10). internal differentiates are restricted to the southern dyke branch between sydtungegletscher and syenitknold where there are two remarkable developments (‘pods’) of layered cumulates in the axial part of the intrusion, comparable to those described above from tuttutooq. the giant dykes close to the inland ice appear to be simple dilational dykes. the combined thickness of the three giant dyke branches at their east-north-eastern-most outcrops is closely similar to that of the two branches at the east-north-east end of tuttutooq and it is probable that the branches continue, beneath the inland ice, to the east coast of greenland. their marginal facies (2–5 m wide) have been generally affected by back-veining from re-melted country-rock granitoids and granitic gneisses with considerable evidence of hybridisation. otherwise, within 10 m of their contacts they tend to be chilled to homogeneous medium-grained dolerites, grading inwards to coarse-grained mesocratic troctolite (fig. 43). plagioclase lamination and modal layering within the coarser-grained axial parts define symmetrically developed synclinal structures in which the dips vary from steep at the margins to horizontal along the dyke axis, much as described by walton (1965) for the giant dyke in johan dahl land. figure 44 shows parallel modal layering in the southern branch across mellemlandet whilst textures in the laminated gabbro are shown in fig. 45. 3838 sydtungegletscher and syenitknold in the southern giant dyke branch, at the sydtungegletscher, well-layered syenogabbros form an axial pod, distinguished from the troctolitic host by their dark colour (due to high content of magnetite) and the presence of prismatic clinopyroxenes. modal layering shows normal grading with erosional troughs filled with mafic cumulates, providing evidence for vigorous magma flow, reminiscent of features at itillip saqqaa on tuttutooq. approximately 3 km to the west-south-west at syenitknold, immediately east of qooqqup sermia, is one of the most remarkable occurrences of differentiated rocks within the ygdc. the course of the dyke makes an abrupt change so that the outcrop on the map resembles a duck’s head pointing into the glacier. on this analogy, the ‘head’ and ‘body’ of ‘the duck’ are 600 m wide. symmetrical border zones of troctolite, carrying plagioclase megacrysts, enclose a central 300 m of mesocratic augite syenite that is traceable east-north-east to the ‘neck’ for c. 1 km. the syenitknold occurrence has much in common with that at asorutit on tuttutooq. however, at the latter the syenite is silica oversaturated, with filter-pressed, pegmatitic alkali granite residuals whereas the syenitknold syenite contains intercumulus nepheline. coarsegrained veins of nepheline syenite, containing aegirine, late-stage fluorite and calcite, intruding the adjacent troctolite, are undersaturated counterparts to the granites in and around asorutit. the syenite is separated from the troctolite by a rusty-weathering zone several metres broad of ferro-syenogabbro grading to ferrosyenite, the colour being due to the high content of magnetite and fayalitic olivine. although these differ texturally from the iron-rich syenogabbroic zones that separate troctolite from syenite at asorutit, (e.g. in lacking dendritic pyroxenes) they represent analogous phenomena. modal layering in the augite syenite dips symmetrically in towards the axial plane; the layering involves rhythmically developed normally-graded units, 10–15 cm thick, in which the bases are defined by melanocratic layers rich in ferromagnesian minerals, passing up into more leucocratic tops. these graded layers are separated by homogeneous layers c. 0.5–1 m thick (fig. 46). although the gardar intrusions as a whole display a wide variety of layering styles, this style is unique to the syenitknold syenite and closely resembles that of the upper zone ferrogabbros of the skaergaard intrusion (wager & deer 1939). fig. 43. coarse troctolite, syenitknold. diameter of coin 2.5 cm. fig. 44. parallel layering in the southern branch of the giant dyke in mellemlandet. 39 a significant feature at syenitknold is the presence, more or less centrally within the syenite, of a crudely tabular, gabbro inclusion estimated to be 50 m thick and 100 m across (fig. 47). although the upper contact has been eroded, it is surmised that this slab was formerly overlain by the syenite and that it composes a large autolithic inclusion of a distinctive gabbro facies. it differs from the host gabbro in being highly feldspathic, containing an abundance of large (up to 20 cm) anhedral plagioclases surrounded by darker olivines. the gabbroic slab is texturally similar to the roofing facies of the gabbro at narsaq (described above) and is accordingly presumed to have been part of the roofing facies of the giant dyke that detached along a subhorizontal joint before collapsing into the underlying residual magma. the latter, which from the mineralogy of the syenite is inferred to have had a benmoreitic to trachytic composition, would have had a density lower than its already crystallised gabbroic roof. a slab of this is inferred to have peeled loose and sunk, finally coming to rest on the upgrowing syenite cumulus. the situation envisaged is shown diagrammatically in fig. 48. the situation described here is reminiscent of that described from the poe mountain anorthosite, labrador, where anorthositic/ leucogabbroic blocks that may have come from a noweroded roof zone sank through resident magma to be arrested on the upgrowing cumulus floor (scoates 2000). fig. 45. lamination in a polished slice of gabbro from mellemlandet. traces of ‘snowflake’ plagioclase growth are discernible in some areas. width of sample 8 cm. fig. 46. layering in the syenitknold syenite. the thin layers, differentiated into melanocratic bases and leucocratic tops, are separated at regular intervals by thicker layers of typical unsorted syenite. 4040 fig. 47. conformable gabbroic layer within layered syenite at syenitknold forms the prominent dark unit in the middle distance. fig. 48. interpretive sections of the geological evolution at syenitknold. a: the giant dyke spread out into a laccolithic or lopolithic body at the unconformity between the julianehåb batholith and the overlying eriksfjord formation. buoyant feldspathic debris from a deep-lying anorthositic protolith accumulated close to the roof. b: the dyke cooled from the sidewalls, while layered synforms grew up from the floor, eventually yielding a trachytic residual magma towards its apex. c: the density of the residual magma was now less than that of the crystallised roof, a portion of which detached to become embedded within upgrowing syenite cumulus. a b c giant dyke magma and troctolitic cumulates buoyant feldspathic debris trachytic residual magma julianehåb batholith eriksfjord formation 41 central complexes and late dykes klokken complex the klokken complex is a stock, approximately oval in plan, (4 km west-north-west–east-south-east × 3 km north-north-west–south-south-east), aligned transverse to the tuttutooq–ilimmaasaq–narsarsuaq lineament and in an isolated position (fig. 49; blaxland & parsons 1975; parsons 1979). it has been u-pb dated at 1166 ± 1.2 ma (table 1), and has a 87sr/86sri of 0.7031 ± 0.0003 (blaxland et al. 1978). the complex has a concentric tripartite structure, the three parts crystallising in sequence inwards. it comprises an outer zone of gabbro (incomplete at the present level of exposure) up to 400 m broad, partially surrounding a broader zone (up to 900 m) of unlaminated syenite. the circular core (c. 2.5 km diameter) of the complex consists of strikingly layered syenites (fig. 50) penetrated in the focal area by a small intrusion of quartz-bearing biotite syenodiorite. marginal gabbro a cross-section and partial cross-section of klokken are shown in fig. 51. the gabbro has near-vertical chilled contacts against the granite-gneiss country rocks. some assimilation between the gabbro and its wall rocks, however, precludes analysis of chilled marginal samples to give approximations of the initial magma composition. evidence from wall-rock pendants suggests that the gabbro body narrows upwards, as depicted in the crosssection that also supposes the complex to be subvolcanic (fig. 51a). the mechanism of emplacement remains enigmatic; xenoliths of gneiss do not accord with ring-faulting and caldera collapse but suggest a process more akin to stoping. apart from gneiss xenoliths the gabbros contain plagioclase megacrysts (high-pressure phenocrysts?) and anorthosite xenoliths that will be described in a later section. within the chilled zone parsons (1979) described development of a wavy pyroxene facies resembling that in the marginal border group of the skaergaard intrusion (wager & deer 1939). the latter facies has recently been interpreted by humphreys & holness (2010) as formed by partial gravitational collapse of crystallising border group cumulates. there are also some subvertical layering features, similar to those of channel-fill structures in sedimentary rocks, indicative of magma flow parallel to the sidewalls together with inward crystallisation. the gabbro, where unaffected by hybridisation, is characterised by stellate clusters (‘snowflakes’) of plagioclase. this feature, very similar to that seen in the ygdc on tuttutooq, is indicative of rapid growth of plagioclase from a melt super-saturated with respect to plagioclase. like the ygdc gabbros, those of klokken are troctolitic, with late and subordinate crystallisation of augitic pyroxene. syenogabbro and unlaminated syenite at high structural levels the marginal gabbro grades into syenogabbro and unlaminated syenite whereas at lower levels there is a distinct break (parsons & brown 1988). modal layering features within it are scarce but those that do occur strike parallel to the contact and are vertical or outward-dipping in the outer section but inwardfig. 49. geological map of the klokken complex. modified from parsons (1979). fig.50 upton-aag 02-10 2013 500 m 600 m 650 m 500 m 100 m 100 m 1 km inland ice gabbro unlaminated syenite layered syenite biotite syenodiorite dip of layering gravel field julianehåb batholith n 45°04´ 60°56´ 4242 fig. 50. oblique aerial photograph of the central part of the klokken complex. the steep escarpments are formed by granular syenite, separated by crumbling, rusty-weathered, laminated syenite. fig. 51. a: cross-section of the klokken complex, postulating an overlying caldera. arrowed flowlines indicate inferred late-stage convection of hydrothermal fluids. modified from parsons & becker (1986). b: more detailed partial section illustrating relationships between country rocks, gabbro, unlaminated syenite and the layered syenites of the central region. modified from parsons & brown (1988). unlaminated syenite layered series sw julianehåb granite ne gabbro late syenodiorite 500 m 1 km s ba n granular syenite layers laminated syenite unlaminated syenite pegmatites present topography gabbro 43 dipping farther in. the unlaminated syenite grades, over 30–100 m, into the central layered series. there is, however, strong cryptic variation within the unlaminated syenites; the fe/mg ratio of the ferromagnesian minerals increases and the feldspars become increasingly capoor and k-rich. the compositions indicate that these progressed in growth from plagioclase through ternary feldspar to sanidine, with subsequent development of exsolution lamellae. feldspar pairs indicate temperature decreases from 950°c in the outer syenogabbros through 910°c in intermediate syenodiorites to 900°c or less in the syenites (parsons & brown 1988). the significance of the break from marginal gabbro to the unlaminated syenite remains debateable. whereas there is no obvious evidence for influx of a more evolved magma, variation in the width of the gabbro sheath and its absence from the southern perimeter suggests that the gabbro underwent thermal or mechanical erosion by new magma. comparable unexplained relationships have been noted above for the ogdc. by contrast, the unlaminated, granular and laminated syenites are all considered to have grown from a single magma chamber. the transition from unlaminated syenite to the central strongly layered syenites is interpreted as marking the change from sidewall cumulate (‘marginal border group’) to lower-angled, centrally directed layered rocks and is analogous to the relationships within parts of the ygdc (e.g. syenitknold). studies of intrusions in which both ‘marginal border group’ and inward-dipping layered cumulates occur, e.g. skaergaard (wager & deer 1939; wager & brown 1968) and parts of the the ygdc, lead to the conclusion that two processes occurred more or less concurrently. crystallisation against the sidewall thermal boundary layer dominated the early stages when heat-loss was higher, and ‘sedimentary’ upgrowth of cumulus talus, descending gravitationally alongside the boundary layer, occurred when thermal insulation was well established. in the first, the growing crystals remained attached to the sidewall whereas in the latter, they were carried down in relatively dense crystal-melt slurries to accumulate above the hypothetical hidden series. at klokken, as in the ygdc and several other gardar cumulate sequences, the factors controlling the relative thicknesses of marginal border groups and central layered series are unknown. extremes range from those that lack any discernible marginal border group (e.g. itillip saqqaa, tuttutooq (upton 1987) and the western stock of the kûngnât complex (upton et al. 2013) to cases where marginal border groups are well developed (e.g. asorutit, krydssø and itillip saqqaa). central layered series the core of the klokken complex is characterised by a layered sequence dipping 30–50° towards a central focus, providing a 650 m thick stratigraphic succession. what makes the central layered series outstanding among all the layered gardar intrusions is the intercalation of more weathering-resistant layers of granular syenite with less resistant and coarser laminated syenites that compose c. 15% of the series (figs 50–52). the laminated syenites show extreme modal layering with inverse grading. here, felsic layers grade upwards into nearly monomineralic pyroxenite layers composed of hedenbergite (up to 90% modal hedenbergite) with interstitial alkali feldspar and titanomagnetite (parsons 1979; fig. 53). furthermore, in places the top (<10 cm) of these inversely graded layers consists almost wholly of fayalite with minor hedenbergite and with interstitial magnetite, amphibole, biotite and alkali feldspar. this rhythm is only completely developed at certain horizons. the mafic/ultramafic upper parts of these layers are orthocumulates in which there is contrast in crystal size between the large tablets of alkali feldspar and the smaller pyroxenes (c. 5 × 1 × 1 mm) and the still smaller fayalites (approximately isometric at c. 1 mm). a further notable feature is the high degree of modal sorting. some pyroxene layers contain >90% hedenbergite but no olivine whereas the olivine-rich layers may contain >90% fayalite. in the inversely graded layers the cumulus phases have slightly more evolved compositions than the same phases in either normal rock or when present as an intercumulus phase in adjacent parts of layers. the layering is attribfig. 52. polished surface (width 10 cm) of laminated syenite from the klokken layered series. 4444 uted to varying degrees of undercooling in a magma in which all phases exhibit a narrow crystallisation interval, and which was subject to rhythmic build-up of pressure followed by sudden pressure releases. crystal accumulation took place under near-stagnant conditions in a thin chamber immediately beneath the roof of the intrusion (parsons 1979). the granular syenites contain sparse phenocrysts of alkali feldspar and the individual units tend to become coarser down the sequence. although there is some feldspar lamination in the lowest units, the granular syenites remain distinct from the enclosing laminated syenite. in the uppermost 100 m of the section, the granular units dominate over the intervening laminated syenites. compositionally, they become increasingly evolved from the highest to the lowest units, showing a regular cryptic variation downwards from more primitive to most evolved while also showing an increase in grain size. the units are interpreted as successive slices of roofing cumulates (i.e. an ‘upper border group’), and the latest (highest) unit may represent a reasonably close approximation to a chilled facies beneath a roof. slabs or slices of units are presumed to have detached serially as planar joints developed during crystallisation, and sank, while retaining mechanical coherence, to become enveloped in the upgrowing cumulus pile that gave rise to the laminated syenites (parsons 1979). in contrast to the strong fractional crystallisation reflected in the granular roofing cumulates, there is only slight cryptic layering in the laminated syenites. it is presumed that there was a sandwich horizon at which well-equilibrated downwardand upward-growing roof and floor sequences met. a hidden layered sequence underlying the lowest laminated syenites is inferred. thus the alternating sheets of the two contrasted syenite types composing the layered central series may be likened to two packs of cards interleaved by a dealer. the granular syenite units were repetitively detached from downgrowing roofing cumulates, whilst the laminated syenites were upgrowing floor cumulates onto which the granular syenite slices came to rest. it may be assumed that the density of the resident magma in the central chamber decreased with time, whereas the density of the roofing rocks with which it was in contact, increased as more and more primitive layers were exposed to it as the detachment continued. that the evolving laminated syenite cumulus remained ductile (or mushy) as the granular syenite slabs settled into it is shown by the development of load structures beneath them (fig. 54). the repetitive detachment and settling of roof cumulates evidenced in the klokken core syenites illustrates the same process described above for the detached roofing block at syenitknold. and, as will be described below, a precisely similar phenomenon took place in the ilímaussaq agpaites. feldspar studies indicate a cooling rate for the syenites of 600 to 500°c in 104 years (brown et al. 1983). although the laminated syenites are deduced to have generally formed as a result of crystals settling in neartranquil magma, evidence that there were sometimes disturbances is provided by some erosional ‘trough and fill’ features and cross-bedding. rare, normally graded layers are believed to have originated from gravity sorting fig. 53. layer in the laminated series, central klokken complex, showing inverse grading from feldspathic base to fayalite-rich top. 45 through current flow. the klokken pyroxenes exhibit a continuum from relatively diopside-rich augites in the gabbros to sodic hedenbergites in the more evolved syenites. the colours change from purplish-brown (in gabbros and unlaminated syenites) to greenish-brown (in upper granulated syenites), apple green (in hedenbergiterich laminated syenites), to deep green sodic hedenbergites (in the lower section of the laminated syenites) and pale green in the late quartz syenites. the more evolved rocks show enrichment in acmite, and in quartz-bearing aplites the pyroxene compositions lie close to end-member acmite (fig. 55). that so many features of klokken are replicated in the ygdc leads to the conclusion that the two intrusions were probably both coeval (table 1) and comagmatic. the principal difference lies in their geometry: the ygdc involved narrow, elongate (dyke-type) magma chambers whereas that at klokken was cylindrical. the compositional ranges, reflected in their respective mineralogies, in both are near-identical. however, whilst the klokken syenites are almost critically saturated with respect to silica (although terminating in silicaoversaturated products), those of the ygdc vary from overto undersaturated in silica. the chronological successions, from chilled gabbros sharing textural and petrographic features, via intermediate rock types within vertically layered border groups to syenite cumulates displaying inwardly dipping layering, are very similar at klokken, krydssø and syenitknold described above. by analogy with evidence from the ygdc, the presence of anorthositic xenoliths in the klokken gabbros may denote proximity to a roof zone in the gabbro. hedenbergitediopside acmitepyroxenes amphiboles fig. 54. load-balls of granular syenite surrounded or penetrated by ‘flames’ from unconsolidated laminated syenite. central klokken complex. hammer c. 30 cm long. fig. 55. pyroxene and amphibole compositions in the klokken complex shown in the acmite–diopside–hedenbergite ternary diagram. modified from parsons (1979). 4646 anorthosite xenoliths and plagioclase megacrysts in the ygdc and klokken gabbros anorthositic xenoliths and plagioclase megacrysts occur in the troctolitic gabbros of both the ygdc and klokken. at syenitknold and asorutit in the ygdc and also at klokken (fig. 56) they are restricted to the outer sheaths of gabbro but are absent from the syenites. at asorutit the gabbro on the north side of the giant dyke contains a crowded array of anorthosite xenoliths up to 100 m across, essentially occupying the full width of the gabbro outcrop (figs 57, 58). the xenolith-bearing gabbro has an apparently fault-bounded contact with the adjacent syenite and it is speculated that this fault downthrows to the nnw, bringing a distinctly high level of the gabbro into juxtaposition with the syenite. this hypothesis would explain the exceptional abundance of anorthositic material, if it is accepted that it is a near-roof facies of the gabbro, crowded with low-density plagioclase-rich material that has floated into place. with increasing distance to the west-south-west along the northern branch of the ygdc on tuttutooq, anorthosite xenoliths and plagioclase megacrysts become scarcer and are absent beyond store pilesø (fig. 5), consistent with the conclusion reached on other grounds that the ‘tugtutôq block’ has been tectonically tilted down towards the ene so that, after erosion, shallower structural levels are seen in the east-north-east and deeper ones in the west-south-west. thus, at both the klokken and tugtutôq complexes, it is concluded that the anorthosite-bearing facies were concentrated towards the top of the intrusion. the anorthositic xenoliths commonly occur together with discrete plagioclase megacrysts up to 0.5 m long. the megacrysts are divisible into cleavage fragments presumed xenocrystal from disintegration of anorthositic autoliths, and subhedral crystals regarded as high-pressure phenocrysts. in the outcrops on the south-eastern side of fabriksbugt at narsaq, the marginal facies of the gabbro at narsaq contains abundant fragments of plagioclase together with some anorthosite xenoliths; similar material is well exposed in road-cuts in the vicinity. anorthosite xenoliths (fig. 40) are abundant around fabriksbugt and extend northwards in decreasing amounts past narsaq and are fig. 56. anorthosite xenolith in the marginal gabbro of the klokken complex, showing anhedral plagioclase, typically with hydrothermally altered mafic minerals in the interstices. pocket knife 10 cm long. fig. 57. laminated anorthosite xenolith on the asorutit peninsula, eastern tuttutooq. the dark intercumulus material is predominantly olivine. sections on scale are 5 cm long. fig. 58. photomicrograph (crossed nicols) of laminated anorthosite from asorutit. it is an orthocumulate comprising cumulus labradorite and intercumulus material mainly consisting of poikilitic olivine. plagioclase crystals up to 12 mm long. 47 absent on the nuugaarmiut peninsula. these observations lead to the conclusion that the xenoliths and megacrysts arrived at their present position through flotation in the troctolitic magma (bridgwater 1967; bridgwater & harry 1968). as related above in the section dealing with giant dykes north of motzfeldt sø, the large gabbroic autolith crammed full of feldspathic debris that occurs centrally within the syenites of the syenitknold is interpreted as derived from a roofing facies within which low-density anorthositic material had accumulated. by the time that a residual low-density body of trachytic magma had been generated beneath the roof, a slab of the roof detached and sank to be arrested within the accreting syenite cumulates. at each of these three localities, the plagioclaserich fragments occur in close proximity to syenite. as the fragments are regarded as indicative of shallow levels in the intrusions, their occurrence is compatible with the conclusion that the syenites were themselves late-stage, shallow-level products generated above upward-grown sequences of gabbroic cumulates. origin of synformal layering in the younger giant dyke complex the layered pods along the ygdc branches exhibit a remarkable variety of layering styles. using an estimated density of 2.8 g/cm3 and a viscosity of 150 poise (g/cm-1/ sec-1) for the initial ygdc magma, mingard (1990) calculated a rayleigh number of c. 1016 and concluded that convection would have been turbulent. however, many of the features observed are best interpreted as products of two-phase (i.e. crystals + melt) convection and, in the more primitive western parts of the intrusion, thermal and compositional convection would have been complementary. why vigorous two-phase convection occurred only at highly localised nodes along the dyke branches remains enigmatic. slurries of crystals + melt, generated in the vicinity of the dyke walls, are postulated to have descended towards a central ‘valley’ within the relatively narrow, deep and elongate magma chambers, the crystals being progressively deposited as slurry velocities decreased. a comparable process is thought to have operated depositing the coarse laminated syenites at klokken. in the more primitive ygdc facies, as seen in western tuttutooq, the relatively fe-rich melt residual from olivine + plagioclase crystallisation would have been denser than the main magma body and would have shown sympathetic downflow. in more evolved melts, in which magnetite had joined the cumulus assemblage, the residual melt would have been more buoyant than the bulk magma and the two effects would then have been antipathetic. however, the evidence throughout all facies of the dyke points to sidewall, two-phase convection in which the crystal-melt slurries were driven by the relatively high modal contents of iron-rich olivines as illustrated diagrammatically in fig. 34. whereas the giant dyke chambers commenced with a tabular, deep and narrow morphology, their geometry would have changed continuously until, in the latest stages (as exemplified by the syenitknold syenites) the residual chamber would have become broad and shallow (fig. 59). similarly (e.g. at klokken), a magma chamber that initially approximated to a deep cylinder with a relatively small diameter would have evolved to a disc-shaped chamber with a quite different aspect ratio. there are indications of a similar morphological evolution of chamber floors at the igdlerfigssalik complex (see below), an evolution that has relevance also at the ilímaussaq complex. a b fig. 59. evolving cumulate morphology in a giant dyke or stock-like intrusion. 4848 mela-aillikites, carbonate-silicate rocks and carbonatites magmatism of an alien character, strongly contrasting with the more voluminous feldspathic gardar igneous suites, occurred intermittently along the southern rift zone and involved ultramafic lamprophyre, carbonatite and carbonate-silicate magmas. such magmas appear to have played a minor role at various times through gardar evolution, and it is postulated that their sporadic recurrence was related to episodic replenishment events as new basaltic magma was introduced into the deep lithosphere, mobilising readily fusible metasomites (upton et al. 2006). these aberrant low-silica magmas gave rise to small hypabyssal intrusions (dykes, sills and plugs) and explosive diatremes. the aillikites contain 20–34 wt% sio2 whilst the more carbonate-rich rocks contain c. 2–10 wt% sio2. this has led to doubt as to whether there is a compositional discontinuity or a continuum, possibly reflecting variable degrees of melting in mantle metasomites (upton & fitton 1985). there are several lines of evidence pointing to a genetic relationship between the ultramafic lamprophyre and carbonate-rich rocks in the gardar province in general, and specifically demonstrated by stewart (1970) and andersen (1997, 2008) for the qassiarsuk volcanic complex. the latter took part in the older gardar activity and has been correlated with the lowest lava member (mussartût member) by andersen (1997). ultramafic aillikitic lavas and sills occur at several horizons within the eriksfjord formation (j.g. larsen 1977; upton et al. 2006). although it is commonly impossible to ascertain their precise chronology, some of the ultramafic magmas appear to have been closely associated with the activity along the younger gardar southern rift. some ultramafic lamprophyre dykes on mellemlandet and in the vicinity of syenitknold cut benmoreite and trachyte dykes, thus establishing their younger gardar provenance. however, other similar dykes north of narsarsuaq in west-southwest mellemlandet are cut by younger gardar doleritic and trachytic dykes confirming the conclusion that such silica-deficient magmas were capable of intrusion over a considerable time period (upton & fitton 1985). the observation that ultramafic dykes are present as integral components of the younger gardar main dyke swarm is itself strongly suggestive that they are all of younger gardar provenance. the age of many of these dykes relative to other gardar intrusions is unknown but their trend and presence within the main and igaliko dyke swarms makes a late gardar age probable. moreover, small carbonatite dykes cut even the youngest parts of the igdlerfigssalik complex which is among the latest major intrusions in the gardar province (table 1). accordingly, as was noted by emeleus & harry (1970), carbonatites occurred throughout a very wide span of gardar time. mela-aillikite intrusions in the narsaq area and on tuttutooq several small intrusions near narsaq comprise ultramafic, silica-deficient alkaline rocks that fall under the definition of ultramafic lamprophyre (rock 1991). they contain >80% (modal) of ferromagnesian silicates and oxides, conferring a colour index of >90. earlier literature referred to them as jacupirangites (ussing 1912) or biotite pyroxenites (upton 1966; upton & thomas 1973) but here, following rock (1986, 1991, 1997) and tappe et al. (2005) they will be described as mela-aillikites, i.e. ultra-potassic (k2o/na2o >3) ultramafic lamprophyres. the field relationships of the mela-aillikites described below leaves little doubt for considering them as postdating ygdc but pre-dating the main dyke swarm. the mela-aillikites are anomalous texturally, mineralogically and geochemically with respect to the majority of gardar igneous rocks. typically they are very fine-grained (50–500 μm) and petrographically very fresh. there are five outcrops of these ultramafic rocks along the west-facing coast of the nugaarmiut peninsula, from the extreme north-west end of the peninsula to c. 1 km from the centre of narsaq (fig. 37). as all five lie approximately at the same stratigraphic level in the host layered gabbros, they may represent protrusions of a conformable, though somewhat irregular, sill-like body (upton & thomas 1973; upton et al. 2006). another, poorly exposed, mela-aillikite occurs on the east side of narsaq township at the water tower. here the mela aillikite is adjacent to a diatreme containing angular clasts of quartzite (presumed eriksfjord formation) and black mafic rock (possibly recrystallised basalt?). yet another (unstudied) occurrence lies on the coast of tunulliarfik, a few kilometres south-east of narsaq. silica activities were too low for feldspars to crystallise whereas perovskite is a common accessory. crystallisation occurred under oxidising conditions ranging from close to the quartz-fayalite-magnetite (qfm) buffer to just below the hematite-magnetite (hm) buffer. in the most highly oxidised facies the clinopyroxenes are bright 49 yellow, pleochroic ferri-diopsides (fig. 60), and compositions of the accompanying olivines approach pure forsterite (up to fo99). apart from other extreme compositions of otherwise common mineral species (very ba-ti-rich biotites and sr-rich kaersutites), the rocks also contain unusual igneous minerals including cuspidine and monticellite. although the principal mineral assemblages appear stable at high temperatures (>600°c), the presence of serpentine, vesuvianite, epidote, chlorite and hydro-garnet indicates that crystallisation persisted to below 400°c (craven 1985; upton et al. 2006). compositionally, the rocks combine high contents of compatible elements (11–24 wt% mgo, 300–1000 ppm ni and 100–1000 ppm cr) with high contents of incompatible elements (craven 1985). such combination of compatible and incompatible elements is a characteristic of other similar rocks such as potassic ultramafic lavas and kimberlites. veining is a characteristic of the mela-aillikites. complexly zoned veins showing bilateral symmetry, up to 10 cm wide, describe curviplanar courses commonly with intersecting flamboyant patterns. the forms of these veins indicate that they were emplaced at relatively high temperatures when their sidewalls were still ductile (fig. 61). it is inferred that the veins mark former conduits for ca-, ba-, sr-, fand co2-rich fluids expelled from volatile-rich magmas during their terminal crystallisation. mela-aillikites are known from two localities on tuttutooq. one is a vertical plug about 80 m in diameter that was intruded up the southern margin of the ygdc, where the latter narrows between the two differentiated pods at asorutit and krydssø (fig. 16). although itself very fine-grained, the mela-aillikite contains corroded olivine megacrysts and small peridotitic xenoliths. meandering late veins, very similar to those of the nûgârmiut intrusions, traverse the plug. the plug has largely obliterated an earlier diatreme containing clasts of partially melted julianehåb granite, quartzite and black mafic or ultramafic rock. these relationships imply that an initial energetic release of gas preceded the ascent of the magma itself. the plug may represent a former conduit supplying a small monogenetic volcano. the other tuttutooq occurrence is a small body of indeterminate size and shape at the intersection of the ogdc and the northern branch of the ygdc. it contains olivine megacrysts, pegmatitic segregations and sparse veins rich in ferrian diopside and phlogopite (craven 1985; upton et al. 2006). at both occurrences, the spatial relationships of the mela-aillikites and the giant dykes suggest that the latter were already in place before the mela-aillikites were intruded. fig. 60. photomicrograph of a highly oxidised pyroxenitic facies in the mela-aillikites. yellow: ferrian diopside. brown to pale brown: phlogopite. black: opaque oxides. colourless: olivine and apatite. field width 2.5 cm. fig. 61. complex, bilaterally symmetrical veins in mela-aillikite on nuugaarmiut peninsula. diameter of coin 2.5 cm. 5050 mantle xenoliths a dyke-like body of aillikitic ultramafic rock, cutting granites on illutalik island c. 7 km south-west of narsaq, is noteworthy for its abundance of ultramafic nodules, interpreted as recrystallised mantle xenoliths (upton 1991; fig. 62). they are rounded, up to 40 cm in diameter, and are almost wholly composed of tremolite and chlorite but with scarce olivine (fo91.5) and chrome-spinels. whilst still identifiable as former peridotites the intense recrystallisation, attributed to deuteric re-equilibration within volatile-rich magma, makes the original petrography debatable. despite the fact that these rocks have been extremely altered they contain cross-cutting veins of glimmerite that may be relatively unchanged. these are principally composed of phlogopite but with small quantities of calcite, apatite, zircon and titanite. they are regarded as providing evidence for k-rich metasomatism that had affected the peridotites prior to their entrainment. diatremes the above-mentioned occurrences in narsaq township and near krydssø on tuttutooq provide evidence for the spatial association of gas-drilled pipes, filled with material that collapsed inwards after venting, and magmatic mela-aillikites. several other diatremes on tuttutooq fig. 62. altered mantle xenoliths in ultramafic lamprophyre host, illutalik. hammer c. 35 cm long fig. 63. diatreme at narsaq. the larger clasts are quartzite (white) and metabasalt (black). height of outcrop c. 1 m. 51 and illutalik are also surmised to be due to degassing of mela-aillkite magmas as are three diatremes that cut the gabbro at narsaq (fig. 37). since these diatremes lie approximately on strike with the five mela-aillikite outcrops at nuugaarmiut, they are also suspected products of mela-aillikite magma degassing. each diatreme is less than 100 m in diameter and is filled with an unsorted assemblage of angular quartzite and metabasalt clasts, presumably derived from formerly overlying eriksfjord formation strata (fig. 63). interstices between the clasts contain calcite, fluorite and gypsum providing support for the concept that the diatremes were generated by highly oxidised, halogen-rich and reactive co2-rich gases or super-critical fluids (upton et al. 2006). several other isolated diatremes in eastern tuttutooq and illutalik may also relate to this phase of magmatism. a diatreme with a carbonated ultramafic matrix cutting the south qôroq complex must be considered as younger gardar (emeleus & harry 1970). other aillikite, carbonate-silicate and carbonatite dykes small (<2 m wide) ene–wsw-trending ultramafic dykes occur in south-eastern tuttutooq and illutalik. the extreme alteration in these is ascribed to volatilerich, low-temperature residual fluids (martin 1985). the lamprophyre dykes in mellemlandet and nunataks to the east-north-east are typically much altered; they are principally composed of fine-grained aggregates of opaque oxides, biotite, carbonate and what may be olivine and pyroxene pseudomorphs. it is likely that late-stage deuteric reactions in the volatile-rich magmas have largely erased early-formed ferromagnesian phases. a swarm of related dykes, trending ene–wsw and exposed along the eastern coast of tunlliarfik fjord north of narsarsuaq, is characterised by very nodular weathering surfaces. the nodules may represent relics of former olivine-rich xenoliths that underwent extensive deuteric recrystallisation. ultramafic lamprophyre dykes with up to 20 wt% mgo are early components of the igaliko dyke swarm (pearce & leng 1996). they consist of approximately equal amounts of diopside and phlogopite, accompanied by opaque oxides, calcite and ferroan pargasite. calcite ocelli occur and one dyke comprises two distinct (streaky) facies, one composed of calcite and the other of carbonate-rich ultramafic lamprophyre with phlogopite and perovskite. a comparable silico-carbonatite dyke composed of alternating streaky layers of calcite and ultramafic rock occurs in the main dyke swarm close to the inland ice (fig. 64). flow-differentiation of materials with contrasted ductility was suggested for the latter (upton et al. 2006). carbonatite (sövite), aillikite and carbonate-silicate dykes occur sparingly among the main dyke swarm eastnorth-east of ilimmaasaq (martin 1985; upton & fitton 1985) but are relatively abundant in the igaliko dyke swarm. these two dyke swarms are described in a later section. the carbonatite dykes, which are restricted to the vicinity of the igaliko syenites, play only a very minor role. whilst the principal carbonate is calcite, other comfig. 64. carbonate-silicate dyke east-northeast of narsarsuaq. white layers are calcitic; dark layers are rich in opaque oxides and silicates. diameter of coin 2.8 cm. 5252 ponents are salite/ferrosalite, phlogopite, apatite, olivine, albite, andradite, perovskite, allanite, bastnaesite, pyrochlore, fluorite and secondary chlorite (pearce 1988). exceptionally, fluorite can compose up to 50% (modal) of the dykes. sr, ba and lree commonly reach wt% concentrations whilst y, nb, zn and th also occur in abundance. fenitisation of their wall-rocks demonstrates alkali loss during their crystallisation (pearce 1988; pearce & leng 1996; coulson et al. 2003). genesis of the ultramafic rocks the case for a genetic relationship between the melaaillikites and carbonatites was made by coulson et al. (2003) and upton et al. (2006), based on a synthesis of sr, nd, c and o isotopic data for these and other gardar carbonatites and lamprophyres (coulson et al. 2003). these authors found no evidence for a compositional gap between the two and concluded that they represent different degrees of melting from the same source. smallfraction, volatile-rich partial melts rising from the asthenosphere became frozen in as lithospheric metasomites before being remobilised during gardar rifting. that a continuum existed from ultramafic aillikites through silico-carbonatites to carbonatites was also suggested for the main swarm dykes in the nunatak regions (upton & fitton 1985). an alternative genetic scheme that the ultramafic lamprophyres and carbonatites are related through liquid immiscibility has been proposed by pearce & leng (1996) as also by andersen (2008) in the case of the older gardar qassiarsuk rocks. the observation that the mela-aillikites at narsaq and on tuttutooq are closely associated with the ygdc invites the suggestion that thermal energy from the latter was responsible for remobilisation of metasomite bodies in the lithospheric mantle, generating the mela-aillikite magmas (martin 1985). the younger gardar aillikite– carbonatite events are inferred to have been short-lived and localised. although they are probably petrogenetically irrelevant to the principal story of the southern rift magmatic system, they afford some insights into the nature of the contemporary lithosphere (coulson et al. 2003; upton et al. 2006). as stated earlier, these ultramafic occurrences have no counterparts in the northern (nunarsuit–isortoq) rift zone. narssaq complex the narssaq complex transects the narssaq gabbro but is cut by the ilímaussaq complex on its eastern flank so that it was intruded during the interval between c. 1163 and c. 1160 ma (table 1). although disturbed by faulting, it 3 km surficial deposits quaternary ketilidian orogen julianehåb batholith eriksfjord formation lavas and sedimentary rocks ilímaussaq intrusion mela-aillikite gardar intrusions narssaq alkali granite narssaq gabbro and associated mafic intrusions narssaq syenite narssaq complex kvanefjeld talut qaqqarsuaq narsaq tunnuliarfik narsaq sund 46° 60°57´ n fig. 65. distribution of syenites and alkali granites of the narssaq complex. note also the intrusions at talut and qaqqarsuaq, which are presumed shallow-level extensions from the narssaq complex. 53 may initially have had an ovoid plan with a diameter of some 10 km (fig. 65). although fig. 65 shows the complex as consisting of two units only: a) quartz syenite and porphyritic pyroxene syenite, and b) alkali granite, this apparent simplicity belies the truth. through a combination of relatively poor outcrop, topographic difficulties and apparent lack of economic resources the narssaq complex has been neglected in comparison with the ilímaussaq complex, its younger neighbour to the east. the bulk of the syenite was presumably emplaced through foundering of the gabbroic lopolith (described above as part of the ygdc) and its overlying cover of eriksfjord formation strata. originally mapped for ggu by j.w. stewart in the 1950s (stewart 1964) and the northern part subsequently mapped in detail by olsen (1977, 1982), the maps and descriptions have not been published. considerable complexity was revealed by olsen’s detailed studies and unpublished map. five intrusive units are distinguished, each with feldspar-phyric margins chilled against the preceding unit. augite syenite i is silica oversaturated. a fine-grained variant contains abundant anorthoclase phenocrysts while another variant is labelled as a black, larvikitic type. augite syenite ii has augite zoned by aegirine-augite and mainly lacks the anorthoclase phenocrysts of augite syenite i. it also contains widespread pegmatites. olsen additionally notes mafic syenite, syenogabbro, leuco-syenodiorite, leucogabbro with anorthosite (presumably as xenoliths) and plagioclase megacrysts. the mountain behind narsaq (qaqqarsuaq; fig. 66) is largely composed of the syenite but is capped by dolerite regarded as part of the former narssaq lopolith. amongst the granite varieties, olsen (1977, 1982) lists microgranite, rhyolite and alkali granite with alkali amphibole. from the author’s own observations some contain aegirine-augite. some mafic layering is described, dipping steeply to ene. in the north-east, low-angled sheets of syenite transgress the metavolcanic and sedimentary strata of the eriksfjord formation which form roof pendants to the intrusion. the present level of dissection is probably close to the roof zone of the complex (emeleus & upton 1976). heterogeneous (streaky) rhyolite crops out on the north-east side of qaqqarsuaq (author’s unpublished field notes) and i. gibson (personal communication, 1974) suggested that this rhyolite could be a caldera-ponded parataxitic ignimbrite. dating of this rhyolite would be desirable to ascertain whether it is part of the narssaq complex or an aberrant component of the eriksfjord formation lavas. on the assumption that it is part of the narssaq complex it would strengthen the case for it all comprising very shallow-level intrusions retaining as well as some extrusive rocks. hydrothermal alteration, pervasive throughout the narssaq complex rocks, may be attributed to fluids expelled during cooling of the ilímaussaq complex that lies 2–3 km to the east. the initial geometry of the complex has been significantly modified by transcurrent faulting. an approximately e–w-trending, left-lateral, transcurrent fault bisects the complex. on the assumption that the fault displaced the ygdc dykes westwards from their position as seen on tuttutooq (at narsaq) to sites now beneath the waters of bredefjord, it is necessary to postulate a displacement of 6–7 km. however, judging from the mapped contacts of the granite (the youngest component of the complex), the movement was much less, possibly only half that distance. accordingly, one may infer that the complex was intruded during an interval of active faulting. the fault fig. 66. view east from narsaq to qaqqarsuaq mountain (685 m). gabbro underlies the foreground and town of narsaq. the qaqqarsuaq mountain is dominantly composed of narssaq syenite and alkali granite, with doleritic sheets. 5454 is likely to have played a significant role in localising the ascent of the narssaq magmas as well as having some influence on the younger ilímaussaq complex to its east. the ilímaussaq complex, described in detail below, contrasts with the narssaq complex in being almost entirely composed of silica-undersaturated syenites. it contains, however, an early intrusion of quartz syenite and highly evolved alkali granite. whereas these oversaturated magmas could have arisen from batches intimately associated with the undersaturated ilímaussaq magmas that had experienced substantial crustal contamination, it may alternatively be speculated that they were residual from the narssaq complex. examples of batches retained within the plumbing systems and making late appearance are known from basaltic volcanoes (e.g. on hawaii) so that the concept of ‘left-over’ narssaq magmas subsequently re-appearing a few kilometres to the east as components in the neighbouring ilímaussaq complex is not wholly inconceivable. south qôroq complex figure 67 is a geological map showing the south qôroq complex adjacent to the north qôroq and igdlerfigssalik complexes. south qôroq is a part of the igaliko syenites that collectively constitute one of the earth’s largest agglomerations of nepheline syenites. rb-sr dating gives south qôroq an age of 1160 ± 8 ma (table 1). it was emplaced across a zone of active faulting and pre-dates the main dyke swarm (described below), sharing these features with the narssaq complex. the south qôroq and narssaq complexes, both emplaced at shallow crustal levels, may possibly mark the sites of two contemporaneous volcanoes, approximately 50 km apart. whereas the narssaq complex transects the lopolithic portion of the ygdc, the south qôroq complex lies several kilometres to the south-east of the giant dykes. it lies across the ese-trending zone of sinistral faults to the north of that affecting the narssaq complex (figs 10, 65). the south qôroq complex, which is partly obscured by the qooroq fjord, was initially mapped by emeleus & harry (1970) and subsequently studied in detail by stephenson (1972, 1974, 1976a). cutting into the eriksfjord formation supracrustal strata, the south qôroq complex reached into the shallow crust and, crystallising largely from low-density phonolitic magmas, it may be a subvolcanic complex. the 5 km i1–7 s1–5 n m x gardar intrusions satellite intrusions motzfeldt complex north qôroq complex ø østfjordsdal complex south qôroq complex members igdlerfigssalik complex early to late members eriksfjord formation julianehåb batholith i2 i5 i4 i4i6 i3i1 s3 s4 s5 s5 s1 s2 s2 i7 ø n m n tu nu l l i a r f ik q oo ro q i ga l iku f jord x x x 45°15´ 61° fig. 67. geological map showing the relationships between the south qôroq, igdlerfigssalik, østfjordsdal, north qôroq, motzfeldt complexes and satellite intrusions. modified from emeleus & harry (1970) and stephenson (1976a), with abbreviated labelling. 55 south qôroq magmas rose through weakened lithosphere adjacent to the older gardar nepheline syenite complexes of motzfeldt and north qôroq (ages in table 1). the close association of all the igaliko syenite complexes is attributed to a lithospheric ‘weak spot’ focussed by the intersection of sinistral faulting and the ene– wsw rifting. the age data in table 1 suggest that a time gap of roughly 100 ma separated the older and younger igaliko complexes during which vigorous plate motion is indicated by palaeomagnetic data (piper 1992, 1995). in view of the close affinities between the older and younger complexes over such a long interval, it is concluded that all shared a similar petrogenesis, presumably from lithospheric rather than asthenospheric sources. although the complex measures 26 km west-northwest–east-south-east and 10 km north-north-east– southsouth-west on the geological map it may originally have had a nearly circular plan with a diameter of c. 10 km (stephenson 1976b). according to stephenson (1976b) the crudely elliptical plan can be explained by ductile deformation of the intrusions while they were still hot, by large-scale simple shear. the complex is cut by alkaline dykes of the main swarm (fig. 68), considered in a later section. intrusion of some of these dykes may also have overlapped with episodes of fault motion (stephenson 1976a). the south qôroq complex is predominantly composed of foyaites. the first intrusion (s1 in fig. 67) occupies a small crescentic area in the far south-east. this was followed by three concentric bodies of foyaite (s2, s3 and s5 in figs 68, 69) inferred to have been sequential stocks with steep outward-dipping contacts. emplacement was by ring-faulting and central subsidence, with younger units engulfing most of their predecessor(s). diffuse or gradational contacts between the three intrusions imply rapid emplacement one after the other. the foci of the successive south qôroq intrusions went through a generalised migration towards the south-east. in view of the high structural level of the complex, an attendant sequence of nested calderas within an overlying volcano may be envisaged. the foyaites are layered cumulates possessing feldspar lamination and modal layfig. 68. nepheline syenite of the south qôroq complex, viewed across tunulliarfik. main swarm dykes are prominent in the middle distance, trending upper right to lower left. 5656 ering with inwardly directed dips. the foyaites were cut by a ring dyke of layered augite syenite (s4), introduced in two pulses, which was itself intruded by a short length of a broad (100 m) syenogabbroic dyke. stephenson (1976a) gives the latter the more precise name of analcime nepheline monzonite. there are also four satellitic intrusions with petrographic characters that encompass most of those forming the main south qôroq complex but which, however, appear to be older than the latter. in view of their high structural level there is the possibility that small phonolitic volcanic cones developed at an early stage, predating growth of the main edifice. the sequence of intrusion s1-s2-s3-s4-s5 deduced by emeleus & harry (1970) was changed to s1-s2-s3s5-s4b by stephenson (1976a). figures 69 and 70 show pyroxene compositions from the south qôroq complex. pyroxenes from s3 are more evolved than those from s5 that, in turn, are more evolved than those from s4b. the olivines in the south qôroq complex are noteworthy for their enhanced ca and mn contents (stephenson 1974). thus their cao contents lie in the 0.2 to 0.4 wt% range, considerably higher than the normal contents for plutonic olivines. as with ca, the mn content of the olivines is abnormally high, averaging 5 wt% mno but with values up to 8.6 wt% mno. the mn concentration increased steadily with fractionation until, in the augite syenite and the foyaites, it became the principal fractionating element in the olivine while fe2+ decreased. the observed olivine range is from fo36fa62te2 to fo2fa82te16. the appropriate annite-alkali feldspar-magnetite buffer curve cuts across the fayalite-magnetite-quartz (fmq) curve so that, in the later stages, the magma following this curve had a fo2 greater than that for fmq at any given temperature. when the two curves crossed, the olivine became unstable and disappeared (stephenson 1974). the successive intrusive units show a compositional trend towards increasingly less evolved compositions. this is regarded as indicative of their having been tapped from progressively deeper levels of a compositionally stratified magma chamber (stephenson 1976a). accordingly, a magma chamber may be envisaged in which highly fractionated phonolitic magma at the top was underlain by silica-undersaturated benmoreitic magma passing down to mugearitic to hawaiitic magma at still lower levels. hedenbergitediopside acmite pegmatite and secondary acmite essexite s4 b ( r) microsye nit e s4b s4 a s5 s2 s3 hedenbergitediopside acmite normal syenite pyroxenes recrystallised pyroxenes secondary acmite syenogabbro pyroxenes pegmatite pyroxenes fig. 69. compositional variations in clinopyroxene in different intrusive members of the south qôroq complex. modified from stephenson (1972) where the prefix ‘ss’ was used instead of ‘s’. fig. 70. pyroxene compositions in the south qôroq complex. modified from stephenson (1972). 57 post-ygdc dyke swarms main dyke swarm the lithospheric extension, and presumed attenuation, along the younger gardar southern rift persisted without abeyance beyond the ygdc event, but at a decreasing rate. the evidence is provided by a remarkable dyke swarm. this, the main dyke swarm, is concentrated along an ene–wsw-trending zone, approximately 10 km wide, that can be traced more than 120 km from the inland ice into the tuttutooq archipelago. although the component dykes never attained widths comparable to those of their giant dyke predecessors, some are up to 30 m wide. there is, however, a generalised decrease in size with increasing youth while the dyke compositions tended to become increasingly more evolved with time (martin 1985; upton et al. 1990). west of ilimmaasaq the swarm exhibits a compositional spectrum from trachybasaltic (hawaiitic/mugearitic) via benmoreitic/trachytic to quartz trachytic, comenditic and (rarely) phonolitic (upton 1964a; macdonald 1969; martin 1985; upton et al. 1990; pearce 1988). throughout the tuttutooq region both the abundance of dykes and their widths diminish notably in the westernmost half of the archipelago (fig. 71), suggesting that their magma sources lay towards the east-northeast. the swarm occupies more or less the same zone as that occupied by the giant dykes and is inferred to have been intruded along the axial part of the southern rift. whereas the giant dykes are unique in size, morphology and composition, the main dyke swarm also represents an outstanding phenomenon lacking any obvious analogue. it has no counterpart in the younger gardar of the nunarsuit–isortoq region and although there are alkaline dykes in the older gardar e.g. in the grønnedal-íka district (emeleus 1964), these compare neither in width nor extent with the main swarm dykes nor with its compositional range. whilst there are many narrow dykes, broader ones with widths >5 m are common, some up to 30 m. distinctive individuals can be followed laterally for up to 40 km. in brief, whilst less spectacular than the gardar plutons and giant dykes, the main dyke swarm represents a major, voluminous, influx of alkaline magmas and, bearing in mind the relatively shallow depth of erosion, it may be suspected that it includes dykes that fed fissure eruptions. some seventy main swarm dykes were recorded in a traverse across the swarm along the east coast of tuttutooq. together with the four giant dyke branches (aggregate width c. 1500 m) and assuming only a 2 m average for the main swarm dykes, this indicates a total c. 1650 m of dyke within a c. 7500 m traverse i.e. a basement dilation here of c. 28%. most of the main swarm dykes are silica saturated to oversaturated and only a small proportion are silica undersaturated (macdonald 1969, 1970; martin 1985; winther 1992). the wide compositional spectrum is ascribed to fractional crystallisation of feldspar, olivine, clinopyroxene ± titanomagnetite and apatite. with evo10 km 46°30´ 60°45´ n fig. 71. map of the post-ygdc main dyke swarm on tuttutooq (dark grey). the tugtutôq central complex is shown in red. 5858 lution from hawaiite to mugearite, the residual magmas were increasingly directed into a residual system, attaining the composition of trachyte approximating to the low-temperature minimum on the albite–orthoclase join. from there the typical trend was towards the alkali rhyolite cotectic and, as emphasized by macdonald (1969), a bundle of affiliated lines of descent was involved rather than a single liquid line of descent. the dyke rocks are typically fineto medium-grained so that whole-rock analyses are taken to approximate the magma compositions. however, with rising silica contents, devitrification textures become increasingly common, and it is known that devitrification of alkali rhyolite glass is accompanied by significant loss of alkalis (especially na) and trace elements. consequently the measured wholerock compositions, particularly of the quartz trachytes/ microsyenites and rhyolites/microgranites, must deviate significantly from the original melt compositions (macdonald 1969; macdonald & edge 1970). the dominance of plagioclase fractionation in the more primitive magmas resulted in feand ti-enrichment, peaking at the stage when mgo had fallen to c. 4 wt% (fig. 72). total fe (as fe2o3) declined from c. 17 wt% in the hawaiites to c. 4.5 wt% in the rhyolites (martin 1985). phosphorus reached a maximum (c. 2.5 wt% p2o5) at the same stage as tio2, reinforcing the conclusion from ygdc studies that titanomagnetite and apatite commenced crystallisation at essentially the same temperature. the highest concentration of ba was reached when mgo had been reduced to c. 2 mgo wt%, approximately at the stage when plagioclase gave way to monoclinic (high-temperature) ternary feldspar. whilst silica-oversaturated (micro-quartz syenitic) dykes extend through the whole length of the southern rift system, rhyolitic dykes are principally confined to the narsaq–tuttutooq sector. conversely, whereas phonolitic dykes are very scarce in the latter, they play a major role in the igaliko dyke swarm further south and east (see below), in the vicinity of the igaliko syenites. big feldspar dykes very distinctive dykes were intruded early in the history of the main dyke swarm. these dykes, which also participate in the igaliko swarm (described below), are characterised by their content of large feldspars (megacrysts) and anorthositic xenoliths (fig. 73). such dykes, known as ‘big feldspar dykes’ (bfds), are not confined to the southern rift but are also widely distributed across the northern (isortoq–nunarsuit) rift as well as still further north into the border zone of the archaean craton. the megacrysts are rarely greater than 50 cm in length, but typically are <10 cm. allaart (1969), however, quotes a size up to 2 m for feldspars from a bfd south-east of the ilímaussaq complex. megacryst compositions range from labradorite to calcic oligoclase and anorthoclase (bridgwater 1967; allaart 1969; winther 1992). bfds commonly exceed 5 m in width and can reach 30 m. since individual dykes can be traced for tens of kilometres, the volume of magma involved was very substantial. these dykes were described in detail by bridgwater (1967), bridgwater & harry (1968) and winther (1992). more recently, detailed investigations were made into bfds of the isortoq area by halama et al. (2002). 1 3 75 mgo% 3 2 1 tio2% fig. 72. whole-rock tio2 vs. mgo in dykes from the nunatak area north-east of motzfeldt sø. fig. 73. big feldspar dyke; main swarm. island south-east of tuttutooq. diameter of coin 2.5 cm. 59 the morphology of the feldspar megacrysts is variable: some are interpreted as corroded high-pressure phenocrysts whereas others are angular cleavage fragments presumed to be derivatives of disintegrating anorthosite masses (bridgwater & harry 1968; halama et al. 2002; fig. 73). the megacrysts and xenoliths are typically confined to the central parts of the dykes, with the outer zones free from, or poor in, megacrystic material. although the relatively fine-grained matrices of the central (inclusion-rich) parts are hawaiitic to mugearitic (containing 55 ± 5 wt% sio2), the marginal facies are distinctly more evolved (benmoreitic to trachytic) and the rocks are described as trachydoleritic and quartz microsyenitic, respectively (bridgwater 1967; bridgwater & harry 1968). however, the terms tephrite, shoshonite and latite are employed by winther (1992) for some of the dykes, emphasising their potassic nature. the margins contain alkali feldspar phenocrysts but are typically devoid of both megacrysts and xenoliths (fig. 74). the phenocryst assemblage in the trachydoleritic central facies comprises plagioclase, olivine (usually pseudomorphed), magnetite and apatite. augite phenocrysts first appear in the mugearite range when the mgo content is down to between 3.5 and 3.0 wt%. from experimental studies on chilled younger giant dyke rocks, the liquid-olivine-plagioclase-clinopyroxene cotectic corresponding with this petrography is attained at a temperature of 1060 ± 15°c at 1kb (upton 1971). the trachydolerite matrices consist of feldspar (zoned from oligoclase to microperthitic alkali feldspar), olivine, magnetite, clinopyroxene, apatite ± hornblende. s and cu reach maxima at c. 4 wt% mgo, inferred to mark the stage at which an immiscible cu-bearing sulphide separated (martin 1985). the widths of evolved marginal facies relative to the more primitive central facies can vary along a single dyke. in the case of one exceptionally wide (20 m) dyke traceable from west-south-west to east-north-east through most of the tuttutooq archipelago, the marginal facies (porphyritic trachyte) expands, from c. 1 m, at the expense of the big-feldspar-bearing trachydoleritic centre, until it occupies the entire width of the dyke. although there is commonly gradation between the two facies, the compositional distinction is abrupt in some instances. thus, in some dykes veinlets of trachydolerite transgress the microsyenite indicating that the latter was solid when the trachydolerite magma remained fluid. the microsyenitic and trachydoleritic facies of the bfds are regarded as cogenetic and their relationship is taken as indicative that they were derived from a compositionally stratified parental magma body in which the more evolved magma overlay the less evolved. during crustal dilation, the benmoreitic-trachytic magma (yielding microsyenite) ascended first, followed, after a variable time interval, by the mugearitic-hawaiitic magma (yielding trachydolerite) that exploited the mechanically weak, still hot, median plane of its predecessor as dilation progressed. the observation that the central facies is generally many times broader than the sum of the border facies may imply that fissure opening began slowly but accelerated as the hotter, inclusion-rich, mafic magma intruded. the bfds contributed to both the main and igaliko dyke swarms and their widespread and highly distinctive characters are suggestive of intrusive phases occurring at a time when very similar compositionally stratified parental magmas bodies existed at depth. whilst it is not claimed that all bfds were synchronous, their features are so idiosyncratic as to make it likely that they were products of a single phase in gardar magmatism. whether or not the hypothesised stratified chambers extended over a great area embracing both the northern (nunarsuit–isortoq) and the southern rift zones remains a moot point. the fact that they brought up copious quantities of xenolithic anorthosite and feldspar megacrysts shows that, like the younger giant dyke magma, the bfd magmas ascended from beneath a solid anorthositic protolith as well as from beneath a level at which large feldspars were growing. study of the megacrysts points to their having experienced complex histories prior to entrainment (winther 1992). the interpretation of some as high-pressure phe~ 1 km 20m ~ 15 km dolerite with anorthosite xenoliths trachydolerite with alkali feldspar phenocrysts trachydolerite with plagioclase megacrysts microsyenite fig. 74. schematic relationships within a big feldspar dyke. modified from bridgwater & harry (1968). 6060 nocrysts augments the conclusions that the anorthosite was comagmatic with the dyke magmas, and that large feldspars were still crystallising at the time that bfds were being emplaced. it is of particular interest that the megacryst population includes not only labradorites but, as noted above, also more sodic plagioclases and anorthoclases (allaart 1969; winther 1992). since no xenoliths composed of such more evolved feldspars have been described, this implies that the latter were high-pressure phenocrysts or primocrysts that became entrained by hawaiitic-mugearitic magma as it ascended rapidly through relatively evolved crystal-melt slurries or mushes which themselves underlay consolidated anorthosite. the rarity of xenoliths and megacrysts in the initial differentiated magma fractions suggests that their densities were greater than that of the magma and inhibited their entrainment. the case has previously been made in the earlier section discussing the more primitive ygdc initial magma, that the feldspathic cargo was buoyant and, hence, capable of flotation. the remarkable abundance of xenoliths + megacrysts (up to 80% (modal); bridgwater & harry 1968) in the bfd trachydolerites may be construed as evidence that there was broad equivalence between the densities of the host trachydolerite (hawaiite/mugearite) magma and its xenolith/megacryst inclusions. it is of interest that granitoid country-rock xenoliths occur in neither facies, probably as a result of being too dense. the impression gained from these remarkable bfds is that even small density differences between melts on the one hand and solid materials (megacrysts and xenoliths) on the other could be of critical importance in controlling whether the latter were entrained or not. such discrimination may also provide further evidence for the low viscosities of the melts. the fact that the magmas could acquire such large quantities of solid labradorite crystals in alkali basalt relatively calcic basalt mafic minerals and bytownite granular anorthosite (an60) high concentration of volatiles layered sequence of bytownite, hypersthene and olivine 2 3 te m pe ra tu re a n d p re ss u re g ra d ie n t g ra d ed m ag m a c o lu m n major syenites anorthosite fragments from roof sink in syenitic magmatranchydolerite with inclusions syenite without inclusions volatile-rich syenite magma s.g. <2.63 basic dyke derived from lower level in chamber andesine megacrysts labradorite megacrysts rare fragments caught up in gardar basic dykes 1 4 5 volatiles and alkalies basaltic magma s.g. c. 2.65 ? (an50–57) megacrysts in intermediate magma s.g. c. 2.63 alkali basalt basalt fig. 75. stages of magmatic evolution in the gardar province, modified from bridgwater (1967). 1: a magma chamber forms. 2: the chamber produces layered floor cumulates and plagioclase flotation cumulates (an60). 3: density of magma lowers and more sodic plagioclase crystallises. 4: compositional stratification forms in in the main magma chamber. the uppermost, more buoyant magma breaks through the anorthositic flotation cumulate. 5: larger volumes of low-density magma (interpreted in the present text to be benmoreitic) ascend to initiate the large gardar alkaline centres. s.g. = specific gravity. 61 detritus and still be capable of intrusion is itself suggestive of low viscosities. bridgwater (1967) proposed that the underlying stratified magma was generated through a process of liquid fractionation in which alkalis and volatiles migrated to and accumulated at the top of the magma chamber. his genetic model is encapsulated in a cartoon, which in principle remains acceptable (fig. 75). the development of stratified magma chambers, analogous to that postulated to explain the relationships in the bfds, is thought to have been responsible for the sequence of intrusions in the south qôroq complex (see above). such stratified magmas may have been generated repeatedly throughout gardar times as exemplified by the kûngnât complex in the older gardar (upton et al. 2013). it is probable that, with passage of time, the volume and depth of the salic upper component progressively increased, thereby reducing the chances for the underlying mafic magma to ascend. salic dykes of the main dyke swarm the differentiated dykes of the main dyke swarm are generally younger than the bfds and most of the cross-cutting dykes are more evolved than those they cut, pointing to the probability that parental magma chambers at depth were undergoing progressive fractional crystallisation (martin 1985; upton et al. 1990). benmoreite dykes, although typically <15 m wide, can attain widths up to 25 m. they are compositionally so similar to the marginal facies of the bfds that they probably arose through the subsequent selective tapping of the uppermost layer of the hypothesised stratified magma body, a layer that may well have grown in volume with the passage of time. feldspar phenocrysts, generally displaying strong zonation, in the benmoreitic dykes have the rhomboid morphology of ternary feldspars and in some instances exhibit the characteristic (100) cross-hatched polysynthetic twinning of anorthoclase (bondam 1955; upton 1964a; fig. 76). phenocrysts of (pseudomorphed) olivine, ferrosalite, magnetite and apatite are typically present. there is close compositional correspondence between the benmoreitic dykes (54–56 wt % sio2 and 1.4–1.7 wt% mgo) and the chilled marginal facies of the ilímaussaq augite syenite (upton & emeleus 1987). however, while the latter is just silica undersaturated, the majority of the benmoreitic dykes are silica oversaturated. magmas crystallising to the quartz trachytes and alkali rhyolites are inferred to have been intimately related to the parent magmas for the assorutit syenite, the narssaq complex and the tugtutôq central complex (described below). the microperthitic, tabular, alkali feldspar phenocrysts of the trachytes and quartz trachytes are presumed to have crystallised as sanidines that subsequently underwent ordering and exsolution. the matrices, often showing well-defined trachytoid textures, comprise alkali feldspar, quartz, biotite, amphibole, aegirine-augite, haematite and primary calcite. with increasing differentiation the amphiboles vary from hastingsite to arfvedsonite. aegirine-augite is present in the more evolved trachytes, as is calcite. whereas devitrification textures are seen only in the fast-chilled margins of the benmoreite and trachyte dykes, the majority of the rhyolitic (comendite or ‘quartzfeldspar porphyry’) dykes represent devitrified glasses (figs 77, 78). the devitrification products are spherical or polygonal spherulites, but in some dykes devitrification led to patchy or finely laminated (flow-banded) rocks coloured deep blue or green according to whether the ferromagnesian component is arfvedsonite or aegirine (fig. 78). buff colours may signify crystalline cores in some dykes. most of the rhyolitic dykes are <5 m wide. phenocrysts in these dykes are hedenbergite, low-quartz paramorphs after high-quartz, and former sanidine, commonly partially exsolved. other components include arfvedsonite, aegirine, astrophyllite, zircon and fluorite (macdonald 1969; martin 1985). the strongly alkaline character of these silicic dykes is typical of comendites. they are rich in incompatible trace elements and zr increases in the main dyke swarm from <200 ppm in the trachydolerfig. 76. rhomb-porphyry-textured benmoreite dyke, eastern tuttutooq. polished slab. diameter of coin 2.5 cm. 6262 ites to >3000 ppm in the comendites. zr, unlike some of the other elements (e.g. li and ga), appears not to have been expelled during devitrification (macdonald & edge 1970; macdonald & parker 1970). the compositions of the salic dykes plot to the peralkaline side of the thermal divide in the system sio2– al2o3–(na2o + k2o (macdonald 1969). there is a gradual increase in na2o/(na2o + k2o) from hastingsite microsyenites to the more evolved arfvedsonite microsyenites. beyond that point, however, there is a regular decrease of alkalis in the more siliceous dykes, ascribed to preferential loss of na2o in a fugitive fluid phase rich in halides and water. there are several occurrences of composite dykes with mafic margins and rhyolitic cores. it is speculated that these arose from underlying chambers in which rhyolite overlay mafic magma and that, as lithospheric pull-apart proceeded, the more mobile deeper magma was drawn up through the viscous rhyolite layer to intrude as a dyke in the overlying country rocks. the hot axial plane of this basaltic pathfinder then lubricated the previously passive rhyolitic magma to permit its intrusion as the younger central component of the composite dyke (e.g. meade et al. 2009; macdonald et al. 2009, 2010). the widest (20– 30 m) and most extensive composite dyke in the tuttutooq region lies close to the northern giant dyke branch and can be traced for 40 km. one of the most extreme compositions in the main dyke swarm is a pantelleritic trachyte dyke on the island of illutalik, off the south-eastern coast of tuttutooq (fig. 10). this 20 m wide, partially devitrified, dyke is remarkable for its conspicuous phenocrysts of narsarsukite (fig. 79). the associated mineral assemblage includes albite, aegirine, zincian nordite, emeleusite, pectolite and unidentified ree-rich silicates (upton et al. 1976, 1978). geochemically the dyke is noteworthy for its low zr/nb ratio, 0.53 (500 ppm zr and 940 ppm nb). igaliko dyke swarm offset from the main dyke swarm but adjacent to it on its southern side is the igaliko dyke swarm (pearce fig. 77. comendite dyke cutting ogdc foyaite, eastern tuttutooq. editor of this volume lotte m. larsen for scale. 63 1988; pearce & leng 1996). although this has also been termed the south-east swarm e.g. by winther (1992), the name igaliko dyke swarm will be used here. this swarm is geographically more restricted, coinciding approximately with the area occupied by the igaliko syenites and traversing the peninsula to their west-south-west, between the tunulliarfik and igaliku fjords (fig. 1). this region is intersected by the same c. e–w-sinistral fault that appears to have controlled intrusion of the narssaq complex 20 to 30 km farther west (figs 2, 80). the variable degrees of offset exhibited by the dykes show that the fault was active at the time the swarm was being emplaced (allaart 1969). a small number of the dykes, however, post-date the faulting and include the remarkable ‘micro-kakortokite’ dyke that will be described in a later section. pearce (1988) pointed out that whereas the geometry of the ygdc branches can be related to shear zones with transtensional extension, the smaller dykes required only simple sinistral shear. the igaliko dykes were emplaced as arrays of en echelon fissures in a sinistral shear regime and the dyke fractures are deduced to have propagated both vertically and horizontally from their source. across a zone c. 15 km broad some 30 km west-south-west of the igdlerfigssalik complex (fig. 2), the abundance of dykes gave rise to a crustal extension of c. 4.5% (allaart 1969). the igaliko swarm dykes cover a wide compositional range but are dominantly silica undersaturated and thus differ from the main dyke swarm. some of the mafic dykes were described by pearce (1988) as lamprophyres; on the total alkali-silica (tas) classification these fall into the fields of tephrites and basanites. alkali lamprophyre (camptonite) dykes within the swarm are distinctive in carrying salite/diopside and kaersutite phenocrysts in a matrix of pyroxene, kaersutite, plagioclase, oxides and a feldspathoid. some bfds showing much the same characteristics as those of the main dyke swarm are also present (allaart 1969). the majority of the dykes, however, are silica undersaturated trachytes and phonolites grading to peralkaline types, to which the mafic dyke magmas are regarded as parental (pearce 1988). according to pearce, up to 70% by volume of the igaliko dyke swarm consists of phonolites with phenocrysts of anorthoclase or sodian sanidine, nepheline and rare salitic pyroxene. research was carried out into the distribution of ree and some other trace elements between phenocrysts and matrices of some of these alkaline dykes to determine partion coefficients between crystals and coexisting liquids (larsen 1979). fig. 78. polished surfaces of devitrified comendites from dykes on tuttutooq. phenocrysts consist of bipyramidal low-quartz (after high-quartz) and perthitic feldspars after sanidine. blue and green colorations in the matrix reflect arfvedsonite and aegirine, respectively. widths about 6 cm. fig. 79. photomicrograph (crossed nicols) of zoned narsarsukite phenocryst in pantelleritic trachyte, illutalik. length of crystal c. 3 mm. 6464 pearce (1988) noted that zr/nb ratios of the mafic dykes provide a discriminant between the main and igaliko swarms: dykes of the former have higher values, averaging 6.4, contrasting with values averaging 3.9 for dykes of the igaliko swarm. thus a whole-rock zr/nb ratio of c. 5.2 effectively distinguishes the two swarms. accordingly the igaliko dyke swarm is relatively enriched in nb and this, in conjunction with its generally more alkaline and silica-undersaturated nature, suggests that its primitive ancestral melts may have been derived from smaller degrees of source rock melting than those of the main dyke swarm. tugtutôq central complex the tugtutôq central complex (tcc) is a small (4.5 × 2.5 km) central complex composed of syenites, quartz syenites and alkali granites (upton 1962; 1964a; upton et al. 1990). it lies astride the ogdc and ygdc and also intersects alkaline dykes of the main dyke swarm (fig. 81). although it is cut by a few ene-trending dykes, the late basic dykes, (described in a subsequent section), it is clearly a very late feature of the southern rift magmatic system, dated at 1156 ± 1.1 ma (table 1). there is, however, no evidence as to whether the complex preor postdates the sinistral faulting that ended the main phase of dyke emplacement. as described previously, the main dyke swarm is considered to represent batches of magma that were episodically released from a deep crustal magma chamber that was undergoing prolonged fractional crystallisation. it is deduced that when extensional rifting was almost finished, a residual volume (>8 km3) of buoyant salic magma ascended by stoping and/or cauldron subsidence. figure 82 is a block diagram illustrating the complex and its principal components. the presence of numerous basalt and quartzite xenoliths within the tcc implies that the eriksford formation formerly extended westwards across tuttutooq and suggests that the complex was emplaced at shallow crustal levels. confirmation of low confining pressures comes from the miarolitic character of many of the tcc rocks. the earliest intrusions of the complex took advantage of the course inflection of the older giant dyke (fig. 5) and gave rise to a small (c. 700 m) plug of porphyritic microsyenite (unit 1a). possibly at much the same time, another intrusion (unit 1b) occurred at a second focal point, 2.5 km to the east-north-east, that transected both n laksenæs fault faults, with sense of displacement gardar intrusions and giant dykes eriksfjord formation inland ice julianehåb batholith 61° 47° n fig. 80. transcurrent faults in the gardar province. those trending e–w to wnw–ese are invariably sinistral. the dextral faults have directions between nnw–sse and ne–sw. 65 the ogdc and the northern branch of the ygdc. this produced a larger body of very similar porphyritic microsyenite. the western microsyenite was apparently then invaded by a coarser syenite that divided the microsyenite into subrounded masses up to 2 m in diameter by a curviplanar network of syenite and quartz syenite veins. the contacts within the plug are diffuse and the veins were intruded when the microsyenite was still hot and ductile, probably above its solidus. the veined microsyenite surrounds a coarser core to this western intrusion. in the eastern centre the microsyenite is present as two annular, steeply dipping sheets (screens) up to 40 m thick that are approximately concentric with the younger components of the eastern centre (fig. 83). the two screens may have originally been contiguous, possibly forming a stock about 2 km across (i.e. over twice the size of that in the western centre). this was later split into the two screens by a ring dyke of alkali granite. as in the western centre, the microsyenite is pervaded by veins of coarser syenite that subdivide it into ovoid masses (fig. 84). in both the western and eastern centres, the veining represents intimate penetration by late-stage fractions of the enclosing magma at temperatures above those permitting brittle fracture. the inner of the two microsyenite screens hosts a large mass (400 × 200 m) of porphyritic country-rock granodiorite as well as smaller biotite-rich mafic xenoliths assumed to have been derived from the eriksford formation basalts and ygdc gabbro. the unit 1 microsyenites are considered to represent magma that crystallised rapidly around the roof and walls of the respective intrusions as a result of loss of heat and volatiles. accordingly they may be considered as analogues of the granular roofing syenites at the klokken complex. four further episodes of roof failure and block subsidence then followed as the eastern centre expanded outwards by stoping and ring-faulting (fig. 85). in so doing it evolved into an ovoid complex 3 km east-north-east– west-south-west and 2.3 km north-north-west–southsouth-east. unit 2 (c. 1200 m diameter) is a heterogeneous quartz syenite that contains an abundance of xenoliths. these include: (a) gabbro from the ygdc, (b) basalt lava, ba1 km unit 6 unit 5 unit 4 tugtutôq central complex unit 2 unit 3 unit 1 younger giant dyke complex older giant dyke complex julianehåb batholith blå månesø 60°51´ 45°21´ n ? ? ?? ? ? ? n 1 km microsyenite fig. 81. geological map of the tugtutôq central complex. simplified from upton (1962). for location see fig. 5. fig. 82. block diagram of the tugtutôq central complex and its relationship to the ogdc and ygdc. the ogdc is farthest from the viewer. modified from upton et al. (1990) fig. 83. map of the unit 1 microsyenites in the tugtutôq central complex. see fig. 81 for location. 6666 saltic agglomerate and quartzite from the eriksfjord formation, (c) granitoids from the julianehåb batholith, (d) clasts derived from the unit 1 microsyenite and (e) clasts from the main swarm dykes. it is hypothesised that a raft of roofing rocks (reminiscent of features in the grønnedal-íka and kûngnât complexes (emeleus 1964; upton et al. 2013) collapsed into the unit 2 chamber. whilst this raft, composed of rocks from above and below the julianehåb granite/eriksfjord formation unconformity, underwent disintegration, it retained its overall stratigraphic integrity. the matrix containing the xenoliths is coarse-grained and, whilst lacking any regular layering features, contains concentrations of olivines and pyroxenes forming mafic/ultramafic schlieren (fig. 86). localised pegmatitic facies of quartz syenite grading to alkali granite add to the heterogeneity. unit 3 ranges from slightly feldspar-phyric quartz syenite to alkali granite and appears to have been intruded with minimal pause after unit 2 as no chilled contacts separate them. unit 3 is largely homogeneous apart from some mafic/ultramafic schlieren like those in unit 2. some schlieren are low-angled with a suggestion of gravity grading. the unit contains scarce clasts of julianehåb granodiorite as well as one great mass (c. 300 × 100 m) of extensively metasomatised ygdc gabbro. where unit 3 is in contact with the unit 1 microsyenite the contact is sharp and dips outwards at 20°. unit 6 units 2–4 unit 1 tugtutôq central complex younger giant dyke complex older giant dyke complex 2 km a b c d n fig. 84. relationship between syenite (pale cream) and microsyenite (dark) in the eastern centre (unit 1) of the tugtutôq central complex. width of sample 12 cm. fig. 85. the inferred intrusive sequence in the tugtutôq central complex. a: giant-dyke disposition predating the tugtutôq central complex. b: emplacement of unit 1 and 2 microsyenites of the western and eastern centres. c: expansion of the eastern centre by units 3, 4 and 5. d: emplacement of unit 6. modified from upton et al. (1990). 67 unit 4 is a coarse-grained hornblende granite, distinguished (commonly with difficulty) from unit 3 in having more quartz, lacking a porphyritic character and containing prominent amphibole. its emplacement is presumed to have involved ring faulting and subsequent cauldron collapse together with the earlier units that are enclosed by it. it appears to form a broad branching sheath around units 1, 2 and 3. numerous thin alkali granite sheets within it dip outwards. unit 4 is largely free from xenoliths but does contain some substantial quartzite xenoliths. chilled margins are again absent and the outward dipping contact with unit 3 is gradational over c. 10 cm suggesting that only a relatively short time interval separated the two magma influxes. external contacts with the julianehåb granitoids are, however, sharp and dip outwards at c. 45°. unit 5 consists of a very narrow alkali granite ring dyke, only a few metres broad. it is traceable for c. 2 km through units 3 and 4 and represents the final stage in the evolution of the eastern centre, possibly marking a culminating caldera collapse within a larger collapse structure associated with the unit 4 emplacement. all four units of the eastern centre appear to have steep outward-dipping contacts as shown in fig. 82. unit 6 is a subcylindrical stock with a diameter of c. 1.5 km that straddles both the ogdc and the northern branch of the ygdc and links the eastern and western centres. this intrusion is unique among the gardar intrusions in consisting almost wholly of alkali feldspar (>95% modal) and can be termed a perthosite. from the roughly crescent-shaped lake (blå månesø; fig. 81) that covers much of its outcrop, it was given the name blå månesø perthosite. whilst the common rock is made up of idiomorphic feldspar crystals c. 2 cm across (fig. 87), the feldspars in randomly distributed pegmatic patches are up to 15 cm long. other finer-grained patches can be termed aplitic. in brief, the unit 6 perthosites are texturally heterogeneous. mineralogy and geochemistry the principal minerals in the tcc are alkali feldspar, quartz, olivine, clinopyroxene (ferrosalite to aegirineaugite and aegirine) and amphibole (ferrorichterite to arfvedsonite). minor minerals include aenigmatite, biotite, ilmenite and magnetite. the olivine ranges from fo9fa88tp3 to fo0fa95tp5 (absent from unit 6) whilst the pyroxenes extend from ferrosalite di35hd62ac3 to virtually pure aegirine. hedenbergite contents reached hd95 before there was any discernible na enrichment (upton et al. 1990). pyroxene crystallisation was generally terminated by reaction to amphibole due to falling temperature and fo2 and rising ph2o and/or pf, although acmitic pyroxene post-dated the amphibole in some facies. the residual melts in units 3 and 4 became highly peralkaline, and in unit 6 aegirine is the sole pyroxene. accessory minerals (ridolfi et al. 2006a) include apatite, zircon, fluorite, thorite, sphalerite, pyrochlore, astrophyllite, ce-chevkinite, yttro-pyrochlore, zirconolite, xenotime, ceriobetafite and ferropyrochlore. the different units of the tcc are principally distinguishable through their textural and modal differences. fig. 86. photomicrograph of fayalite-hedenbergite-magnetite-rich cumulate in unit 3, tugtutôq central complex. olivine partially replaced by opaque oxides. intercumulus alkali feldspar. fig. 87. typical exposure of the blå månesø (unit 6) perthosites. hammer c. 40 cm long. 6868 cathodo-luminescence studies of the blå månesø feldspars demonstrated interconnected pores (a microporosity of 4.1%) that permitted passage of metasomatic fluids with consequent large-scale interaction. the fluids flowed along grain boundaries via the micro-pore network (finch & walker 1991). the relative abundance of fluorite in units 2 to 6 points to concentration of fluorine in the magmas, and the marked metasomatism that affected the inclusions is attributed to the reactivity of halogenated melts or fluids. na-rich aqueous fluids, inferred to have persisted to subsolidus temperatures, caused secondary alteration of the perthites to clays (ridolfi et al. 2006a). primary carbonate is only rarely seen, occurring interstitially in some syenites of the eastern centre (upton 1964a). however, there is evidence that carbonatitic fluids, rich in f, na, ca, p and lanthanides, permeated the rocks, leaving their mark in the form of distinct post-magmatic textures and mineralogies. these fluids also generated albite, fluorite, ce-monazite and almost pure ce-bastnaesite within rock fractures and vugs. evidence for selective interaction between the early mineral phases and late-stage lree-rich fluids comes from the scattered patterns shown by whole-rock geochemical plots of zr vs. lanthanides. the nature of the post-magmatic phases implies that the hydrothermal fluids were enriched in na, ca, p, lree, f and co2. fluid interaction took place at temperatures <550°c. activity by co2-rich fluids followed and, at lower temperatures (150–250°c), by h2o-rich fluids (ridolfi et al. 2006a). whole-rock compositions from the tcc contrast with those of the main swarm dykes with comparable silica contents (58–74 wt% sio2) in having more al2o3, k2o and na2o and less mgo, fe2o3 (total iron), mno and tio2 (fig. 88). whereas the compositions of the more rapidly crystallised main swarm rocks may be taken to roughly equate with those of their magmas, the tcc rock compositions are believed to deviate significantly from their melts because of differential loss of high-density (fe-rich) minerals that left the residual magma correspondingly richer in feldspar components. this process was most profound in the residual magmas of unit 6. the consequent perthosites are therefore regarded not as alkali feldspar flotation cumulates, but as products of residual magma following gravitational depletion in ferromagnesian minerals. the composition of this residual magma approximated closely to the minimum melting point composition on the albite-orthoclase join (upton et al. 1990). lower zr/nb ratios in the tcc relative to main swarm salic dykes may also be due to selective loss of zircon through crystal settling. petrogenesis the earliest components of the tcc (unit 1) appear to be more primitive (resembling the preceding benmoreite dykes) than those of units 2, 3 and 4. the interstitial mineral assemblage of the unit 6 perthosites demonstrates that the magma was more highly fractionated than those of the earlier tcc units and supports the thesis that the magmas became increasingly evolved with time. the conclusion is that benmoreitic magma at depth was steadily evolving towards a peralkaline quartz trachyte composition (upton et al. 1990), i.e. essentially what was concluded in the case of the main dyke swarm. clearly it was not a continuation of the latter but rather that it was a localised repetition beyond the time when lithospheric attenuation had virtually ceased. the puklen complex in the northern (nunarsuitisortoq) rift shares numerous features with the tcc. it is similar in size and was also intruded across a gabbroic giant dyke (pulvertaft 1961, 1965; parsons 1972). both complexes consist of syenites, quartz syenites and peralkaline granites. an isotopic study of puklen suggested that whereas the syenites may be regarded as mantle derivatives variously modified by assimilation of upper crustal materials, the o isotope data for the granites imply either a different source or a different crustal contaminant (marks et al. 2003). these conclusions may also be applicable to the tcc. n a 2 o + k 2o % 12 10 8 6 4 sio2 % tugtutôq central complex main swarm dykes 42 46 50 54 58 62 66 70 74 fig. 88. alkali/silica plot of samples from the tugtutôq central complex and the main swarm dykes. star symbol at c. 67 wt.% sio2 represents minimum melting point composition on the albite-orthoclase join. modified from upton et al. (1990). 69 comparable phenomena to those of the tcc and puklen are known from syenite autoliths within the kilombe volcano in the kenya rift. kilombe may provide a modern analogue for the volcano inferred to have overlain the tcc. it was deduced from kilombe that carbonatitic fluids rich in f, na and ree percolated the subvolcanic system, interacting with the syenites at the thermal boundary layers of the magma chamber during and after their crystallisation (ridolfi et al. 2006a, 2006b). if the tcc was overlain by a central volcano comparable to kilombe, this would have been largely composed of trachytic and comenditic extrusives. the western centre may have been crowned with a small trachytic cone whereas the eastern centre probably underlay a nested set of concentric calderas. the unit 6 magmas may have underlain a culminating caldera, developed by late collapse of a peralkaline salic volcano. on the basis of macdonald and smith’s (1968) hypothesis relating the area of calderas to the volume of extrusives, the tcc may have erupted <6 km3 of magma. late basic dykes whereas intrusion of the main dyke swarm was essentially ended before the tugtutôq central complex was emplaced, there was some small-scale revival of dyke activity in post-tcc times (martin 1985) that produced sparse, thin (<2 m) dykes with the same regional ene– wsw trend as the rift as a whole. a few are trachytic but other basic dykes are distinctive in that they are typically flow-banded and contain megacrysts and in some cases xenoliths and also generally contain ocelli (fig. 89). the volatile-rich nature, ocelli, megacrysts and phenocrysts indicate a lamprophyric and, more specifically, a camptonitic character. similar, strongly altered, camptonite dykes cut the kvanefjeld area of the ilímaussaq complex (sørensen et al. 1974, larsen 2006). the ene–wsw trend of the camptonites strengthens the case for their acceptance as gardar intrusions and they are provisionally accepted as of younger gardar age although they have not been dated. a very fresh, nne-striking monchiquite dyke that cuts the ilímaussaq intrusion on kvanefjeld (larsen 2006) has been rb-sr dated to 1134 ± 17 ma (table 1). the unaltered character of this dyke suggests it was emplaced after the alteration of the camptonites, also strengthening a gardar age for these. the camptonite dykes contain phenocrysts of plagioclase and titan-augite. the ocelli, generally occupied by chlorite and calcite ± albite and epidote, compose up to 15% by volume of the dykes. single feldspar megacrysts (up to 5 cm large) occur as do composite aggregates comprising feldspar, magnetite and ferromagnesian minerals. the feldspars are normally-zoned, an66–33. the megacryst assemblages include amphibole (potassian kaersutite) and titanium-rich biotite as well as clinopyroxene, magnetite and apatite. it is the concentration of megacrysts and phenocrysts into subparallel layers that gives rise to the flow-banding of the dykes. compositionally the late basic dykes resemble the main swarm hawaiites but differ in being olivine-free, more volatile-enriched and silica undersaturated. major element ratios typifying both dyke sets are similar and both are typical of gardar mafic rocks in general (i.e. with high al/ca and low mg/fe). these late dykes, however, have lower al/ca but significantly higher mg/(mg+fe), fig. 89. late basic dyke, nasaasarli, tuttutooq. hammer c. 50 cm long. 7070 ni, cr and v than the main swarm hawaiites. the chondrite-normalised incompatible element patterns of the late basic dykes are relatively smooth, much like those of the main swarm hawaiites. they lack any significant eu anomaly and have lan/ybn values between 12.8 and 17.5 (martin 1985). however, relative to main swarm hawaiites, the late basic dykes are less enriched in ree. if they are late components of the southern rift (tuttutooq–ilimmaasaq–narsarsuaq) magmatic system these dykes are important in signalling a renewed mantle melting episode, albeit on a minor scale. ilímaussaq complex the original description of this extraordinary intrusion was given by n.v. ussing (1912) and over the past forty years a plethora of publications has appeared adding much detailed information. general reviews have been presented by ferguson (1964), larsen & sørensen (1987) and sørensen (2001, 2006). the ilímaussaq complex has long attracted attention for its exotic mineralogy and for the layered syenites. the petrogenesis of its rocks has long been debated and still remains contentious. in the 1950s and 1960s ilímaussaq received close investigation because of its potential as a source of uranium. subsequently its content of rare metals including zirconium, beryllium, niobium, and tantalum brought it to the attention of mining companies. most recently the possibiltalut tuttup attakoorfia naajakasik siorarsuit ilimmaasaq kvanefjeld tupersuatsiaat appat qeqertaasaq nunarsuatsiaq illunnguaq lil l ee lv laksee lv tase q søndre kangerlu arsu k nunasarnaasaq n ar sa q elv nakkaalaaq nunasar naq ki lla va at tunulliarfik laksefjeld kr ing ler ne 3 km surficial deposits quaternary augite syenite ygdc intrusive sheet alkali granite and quartz syenite ketilidian orogen julianehåb batholith eriksfjord formation lavas sandstone arfvedsonite lujavrite lujavrite transition zone mediumto coarse-grained lujavrite aegirine lujavrite kakortokite agpatitic nepheline syenites naujaite sodalite foyaite pulaskite, foyaite marginal pegmatite ilímaussaq intrusion other gardar intrusions narssaq alkali granite narssaq syenite gravel road 45°50´ 60°55´ fig. 90. geological map of the ilímaussaq complex. modified from ferguson (1964) north of tunulliarfik, andersen et al. (1988) south of tunulliarfik, and sørensen (2001). 71 ity of exploiting different rock facies for rare-earth elements has kindled worldwide interest. the complex has a crudely ovoid plan c. 18×8 km, elongate north-west–south-east, transverse to the southern rift zone (figs 90, 91). it has been precisely dated at 1160 ma (table 1). although intruded across the transcurrent fault that previously accompanied and displaced the narssaq complex, it post-dated all significant movements along it. nonetheless, the form of the intrusion appears to have been partially dictated by the fault, particularly the marginal embayment in its north-western sector. stephenson (1976b) suggested that the overall elliptical outline of the complex was a product of strain when it was still ductile, as in the cases of the south qôroq and igdlerfigssalik complexes. planar deformation flaser structures occur in the earliest component (augite syenite) beside narsaq elv (narsaq river) close to the fault extrapolation (hamilton 1964). a photograph in ferguson (1964) portrays stretched naujaite (a rock type described below) that appears to have undergone ductile deformation. although the precise locality is not given, it is from “n.w. of taseq lake” which could put it close to the eastward extrapolation of the transcurrent fault. these deformations in the augite syenite and naujaite suggest that seismic stability had not been entirely achieved at the time the complex was emplaced. on the eastern slopes of kvanefjeld (in the north-western part of the complex) the country rocks are fractured and sheared close to the contact. the volcanic roof in this locality dropped by 300 to 400 m through faulting prior to consolidation of the agpaitic rocks beneath (j.g. larsen 1977). some 12 km to the east-southeast, at nunasarnaq on the north side of tunulliarfik fjord, the eriksfjord formation sandstones and lavas are strongly sheared (sørensen 2006). such shearing is atypical at the margins of the gardar plutons; emplacement by stoping would not entail such deformation and the implication is that faulting had taken place prior to magma emplacement. as with the narsaq complex, intersection of the transcurrent fault and rift axis fissuring is presumed to have provided the potential conduit that was exploited by buoyant magmas. the magmatic focus, however, had now relocated from the narsaq area several kilometres eastwards along the fault zone. a critical hinge-fault, traversing the southern part of the ilímaussaq complex, can be traced east-north-eastwards from the kangerluarsuk fjord and along lakseelv (fig. 90). this fault divides the complex into a southern portion containing the lowest exposures and a larger northern portion that reveals shallower levels in the intrusion (ferguson 1964; bohse et al. 1971; sørensen 2006). the downthrow on the northern side diminishes towards the east-north-east from at least 600 m in kangerluarsuk towards zero as it reaches appat on the southern coast of tunulliafik. the northerly downthrow relates to two (and possibly three) successive movements: pre-ilímaussaq, post-aegirine lujavrite, and possibly postarfvedsonite lujavrite (bohse et al. 1971; sørensen 2006). in the structurally higher northern part the upper part of the complex is preserved, with eriksfjord formation strata as the principal country rocks. the exposures in the deeper southern part below the base of the eriksford formation show the ilímaussaq augite syenite in contact with the julianehåb batholith granitoids (fig. 90). fig. 91. northern part of the ilímaussaq complex viewed across tunulliarfik. 7272 until recently there was general consensus that the complex formed through three successive intrusions but recent work (sørensen et al. 2006) suggests that there were four or possibly more, each of which followed sufficiently quickly to preclude development of well-chilled internal contacts. accordingly the entire ilímaussaq assemblage would have cooled as a single thermal unit. all the components are believed to have originated from a single parental source, with fractional crystallisation governed by low water activity, low silica actvity and low fo2 (engell 1973; larsen 1976, 1977; larsen & sørensen 1987; marks & markl 2001). crystallisation of the entire complex took place over an extended temperature interval of at least 950–450°c (larsen & sørensen 1987) and possibly persisting down to 300°c (marks et al. 2007) with the closing phases being marked by an abundance of pegmatites and hydrothermal veins (engell et al. 1971). in this account it will be assumed that there were three principal intrusions, yielding: 1) augite syenite, 2) alkali granite and quartz syenite, 3) agpaites, but bearing in mind that the actual number of agpaitic influxes remains controversial. agpaite was the name bestowed by ussing (1912, p. 341) on these highly peralkaline rocks. he defined them as follows: “thus if na, k and al are the relative amounts of na, k and al atoms in the rock, the agpaites may be characterized by the equation (na + k)/al ≥ 1.2, whereas in most ordinary nepheline syenites the ratio does not exceed 1.1”. ussing called this ratio the ‘agpaitic index’ but it is more correctly termed ‘the peralkalinity index’ (sørensen 1997). agpaites are peralkaline nepheline syenites containing aegirine, sodic amphibole and/or aenigmatite as well as complex zrand ti-silicates. they are rich in f, cl and h2o (sørensen 1960) and are characterised by exceptionally high contents of zr, hf, nb, ree, u and a host of other highly incompatible elements (bailey et al. 2001). the ilímaussaq agpaites are regarded as the products of extremely fractionated iron-rich phonolitic magmas and include some of the most evolved and incompatible-element-rich rocks on earth. they comprise the rock types sodalite foyaite, naujaite, kakortokite and lujavrite. for several decades after ussing it was considered that a single agpaitic magma body had differentiated into a downgrown roof sequence, a complementary upgrown floor sequence of layered cumulates, and a trapped ‘sandwich horizon’ that crystallised between the two (fig. 92). this hypothesis, proposed by ussing (1912), that an initially homogeneous magma had crystallised as a closed system to produce the contrasting rock types was, for many years, generally accepted (e.g. ferguson 1964, 1970a 1970b; sørensen 1969; engell 1973). studies over the past thirty years, however, have shown that the closed system model for the agpaitic magma is too simplistic although it retains its adherents. whilst belief in the ultimate consanguinity of the agpaitic rocks remains unshaken, at least two magma influxes of peralkaline composition are now proposed (sørensen 2006) and the possibility of multiple replenishments is being considered. there is general consensus that the rocks crystallised at a pressure of c. 1 kb, corresponding to a depth of 2–3 km (j.g. larsen 1974; konnerup madsen et al. 1979; krumrei et al. 2007). heat-flow data suggest that the total thickness of the agpaitic rocks, with their high contents of radioactive elements, should be less than 1 km (sass et al. 1972). gravity and density data gave a best-fit model showing that a heavy body, with density of least 2.9 g/cm3 and vertical boundaries, underlies the complex ew 1000 1500 m 0 500 ??? 2000 m floor series augite syenite alkali granite and quartz syenite naujaite sodalite foyaite pulaskite, foyaite lujavrite kakortokite marginal pegmatite roof seriesmarginal zone rocks fig. 92. simplified section across the ilímaussaq complex. from andersen et al. (1981a). 73 at a depth of 2–5 km (forsberg & rasmussen 1978). as emplacement of the complex must have involved foundering, not only of julianehåb granitoids but of the basaltic eriksfjord formation lavas and gabbros of the ygdc lopolith, this heavy body is likely to include these preilímaussaq country rocks together with any cogenetic high-density cumulates. augite syenite augite syenite is present as a partial shell around the western and southern sides of the complex. the augite syenite probably originally formed a single body almost all of which, except for the remnant shell, foundered within the younger agpaitic magma(s). the focus of the agpaitic activity migrated eastwards by <1 km in relation to its augite syenite predecessor. the augite syenite contacts vary from steep to vertical to outwardly dipping beneath quartzites and an intrusive sheet (older gardar) between tunulliarfik and kangerluarsuk fjords. in the northern part of the complex the augite syenite has a subhorizontal contact with eriksford formation trachytic lava (ferguson 1964). whereas there is consensus that the youngest agpaitic magmas crystallised essentially in a closed system beneath a more or less impervious capping, this is less certain in the case of the augite syenite. the augite syenite is deduced to have been emplaced by block subsidence (sørensen 1978; nielsen & steenfelt 1979). evidence of some piecemeal stoping, however, is provided by clasts of quartzite up to 100 m across derived from the eriksfjord formation, that occur in the augite syenite on the southern shore of kangerluarsuk fjord (ussing 1912; ferguson 1964). since these quartzite xenoliths occur far below the eriksfjord formation – julianehåb granite unconformity, their presence implies that the magma had both low density and low viscosity. the pattern for emplacement among several gardar plutons involved repetitive collapses of slices or rafts of roof rocks (providing temporary ‘floors’ for cumulate deposition) and this process may also have characterised the augite syenite emplacement. on either side of kangerluarsuk fjord the augite syenite has well-chilled margins against the country rock granitoids although on the northern coast there is notable rheomorphism. the augite syenite also shows signs of chilling against eriksfjord formation trachyte in the summit area north and north-east of taseq (fig. 90) where the roof zone is exposed (ferguson 1964). strongly sheared augite syenite xenoliths are found at kvanefjeld in the far north-west of the complex and within lujavrite in a contact breccia (mélange) at the western contact on the northern coast of tunulliarfik. augite syenite is also seen as large xenolithic masses in the kakortokites which, as described below, form the lowest exposed unit in the complex (ferguson 1964; sørensen 1978, 2006; nielsen & steenfelt 1979). tilting of eriksfjord formation strata in towards the intrusion suggests that country rock engulfment during emplacement of the augite syenite and/or the agpaitic magma(s) was accompanied by down-drag of the adjacent crust. however, the inward dips may also relate to very late-stage downsagging of the central part of the complex (ussing 1912; sørensen 2006). fine-grained marginal facies of the augite syenite show the magma to have been a silica-undersaturated benmoreitic magma, closely related to the benmoreites of the igaliko dyke swarm (upton & emeleus 1987). apart from some relic oligoclase (hamilton 1964; larsen 1981), zonation in feldspars from the chilled marginal facies shows the rhomboidal form characteristic of early-formed anorthoclase (upton 1964a; larsen 1981). the chilled marginal feldspars are identical to those of the south qôroq augite syenite s4 (stephenson 1976a; larsen 1981). the feldspars are mostly untwinned cryptoperthite to microperthite, with compositions an20ab76or4 to an3.5ab43.5or53 straddling the oligoclase – ternary feldspar (anorthoclase) – sodic sanidine range. nepheline occurs interstitially. the early ferromagnesian minerals are ferrosalitic pyroxene (100 mg/ (mg+fe2++mn) = 52–21) and olivine (fo17.3-4.0) closely mirroring compositions in the eastern stock at the kûngnât complex (larsen 1976; stephenson & upton 1982). these are accompanied by amphibole (titanian ferroan pargasitic hornblende), titan-biotite, nepheline, magnetite and apatite (larsen 1976, 1981). the ternary feldspar crystallised at c. 1000°c and according to marks & markl (2001) and markl et al. (2001) it was joined by magnetite, olivine and augite within the interval 800– 650°c. however, on the presumption that the parental augite syenite magma was a younger batch from the same source as the preceding ‘rhomb-porphyry’ benmoreite dykes, it is more probable that it already contained phenocrysts of feldspar, olivine, augite, titanomagnetite and apatite at the time of intrusion. in the deeper section of the complex, south of the kangerluarsuk–lakseelv fault, the augite syenite exhibits various forms of modal layering. some gravity-stratified modal layering dips steeply inward (fig. 93). this is 7474 comparable to features in the nunarssuit and kûngnât syenites ascribed to marginal, downflowing slurries of melt and primocrysts that deposited their crystal component as the flow velocity diminished and the sidewalls graded into lower-angled chamber floors (upton et al. 1996). however, repetitive inch-scale isomodal layering is also developed in these southernmost outcrops (ferguson 1964). the inference is that the syenite formed a layered, stock-like body from a magma chamber in which two-phase (i.e. liquid+crystals) convection developed. the older gardar eastern stock of the kûngnât complex appears to provide the closest analogue in the gardar province (upton 1960; upton et al. 2013). alkali granite and quartz syenite two silica-oversaturated sheets cut the augite syenite in the highest parts of the complex. the quartz syenite forms a layer above the pulaskite (described below) and is overlain by alkali granite (ferguson 1964; steenfelt 1981). the age relationships between the quartz syenite and granite are indeterminate and it has been suggested that the quartz syenites are products of interaction between the granite and nepheline syenite magma (sørensen 2006). the alkali granite (fig. 94) is hypersolvus and comprises c. 54% perthite, c. 30% quartz, c. 15% aegirine+ arfvedsonite and c. 1% accessory minerals (hamilton 1964). early, untwinned, alkali feldspar is inferred to have been replaced by microcline perthite and later, granular albite. the feldspars contain a host of minute (<100 μ) aegirine prisms, possibly exsolution products, that confer a green colouration to the rock (fig. 94). arfvedsonite, with some relic aenigmatite cores (larsen 1977) has rims partially replaced by aegirine. the granite magma was highly enriched in incompatible elements (bailey et al. 2001) manifest in the presence of e.g. nazr and na-be silicates (elpidite and epididymite). other minor components include astrophyllite, pyrochlore, leucosphene, fluorite, calcite and zircon (ussing 1912; hamilton 1964). the la/ybn ratio is 9–10 and there is a marked negative eu anomaly. crystallisation occurred at, or above, 750°c but the late growth of aegirine (at the expense of arfvedsonite) took place at, or below, 350°c (markl et al. 2001). fig. 93. modal layering in augite syenite on the south coast of kangerluarsuk. hammer c. 40 cm long. fig. 94. polished slab (11 cm wide) of the ilímaussaq alkali granite. the green colouration is due to micro-inclusions of aegirine in the alkali feldspar. grey areas are quartz; black is mainly arfvedsonite. 75 agpaitic syenites the rocks crystallised from the agpaitic magma comprise a roof series of pulaskite, foyaite, sodalite foyaite and naujaite, a floor series of kakortokites and lujavrites, and a ‘sandwich horizon’ of lujavrites. the total thickness of agpaitic rocks is estimated at c. 1600 m (andersen et al. 1981a; krumrei et al. 2007). by the time the lowest exposed kakortokites were deposited, the roof series was already solidified and c. 800 m thick (bohse & andersen 1981) so that the naujaite had cooled to, or below, its solidus at c. 500°c (andersen et al. 1981a; konnerup-madsen & rose-hansen 1982). subsidence of the chamber floor probably occurred incrementally rather than in a single event while the floor cumulates were being deposited. the occurrence of naujaite and foyaite autoliths at various levels within the kakortokite-lujavite succession suggests episodic roof instability. at nunasarnaq (eastern contact, north coast of tunulliarfik, fig. 90) the magma chamber wall appears to have collapsed, with detachment of large xenoliths (some several hundred metres across) of naujaite and eriksfjord formation basalt into unconsolidated lujavrite (sørensen 2006). and at the southern and eastern contacts vertical fractures present in the augite syenite and adjacent julianehåb granitoids are thought to be related to the ‘rafts’ of these rocks that collapsed into the chamber during emplacement of the lujavrite. at kvanefeld (at the north-west margin of the intrusion) xenoliths of basalt, gabbro, anorthosite, augite syenite, naujaite and alkali syenite lie within the lujavrite (ferguson 1964; sørensen 2006; sørensen et al. 2011). the agpaitic magmas are regarded as residual after very high degrees of fractional crystallisation from an augite syenite parental magma (engell 1973; bailey et al. 2001) and are inferred to have been highly fractionated, iron-rich phonolites. their crystallisation was controlled by low activities of water and silica in conjunction with low oxygen and sulphur fugacities (sørensen et al. 2006). the time span for fractionation within the agpaite magma has been shown by ar data to have been not merely less than 5 ma, but possibly much shorter, probably of the order of 500–800 ka (krumrei et al. 2006). molecular ch4 and other hydrocarbons present in fluid inclusions in the agpaites have generally been regarded as of magmatic origin (konnerup-madsen & rose-hansen 1982; konnerup-madsen et al. 1988; konnerup-madsen 2001). however, this has recently been challenged by laier & nytoft (1995, 2012) who argue that the hydrocarbons contain characteristic biomarkers and the carbon isotope signatures point to an organic origin, probably originating from downward percolation of fluids from much younger mesozoic–cenozoic sediments. although the great bulk of the agpaites consists of cumulates or pegmatites some rocks (e.g. the finer-grained facies of the marginal pegmatite and some of the final lujavrites and the micro-kakortokite dyke, each described below, have bulk compositions thought to approximate to those of melts (larsen & steenfelt 1974; larsen & sørensen 1987; sørensen 2006). roof series pulaskite, foyaite, sodalite foyaite. rocks constituting the roof series of the agpaitic part of the complex are preserved beneath a cover of augite syenite and/or eriksfjord formation lavas. although fig. 92 shows the agpaitic roof to be approximately horizontal, it is distinctly irregular (sørensen 2006). the roof rocks are notably coarsegrained with much conformable pegmatite, attributable to accumulation of volatiles beneath an impermeable roof (ferguson & pulvertaft 1963; ferguson 1964; larsen & sørensen 1987; sørensen 2006). a downward accreting crystallisation front created a roof series comprising a four-member sequence (ussing 1912; ferguson 1964; larsen & sørensen 1987; sørensen 2006). from the top down these are pulaskite, foyaite, sodalite foyaite and naujaite, produced successively from increasingly fractionated melts. thus the sequence has some analogy with the upper border group at skaergaard (wager & brown 1968) as well as with the inferred roof series at klokken. although the pulaskite and foyaite do not themselves qualify as agpaites they are regarded as the earliest (roof) products from the magma body from which the true agpaites crystallised. the sodalite foyaite crops out over a wide area above the naujaite between the tunulliarfik and kangerluarsuk fjords. engell (1973) considered the bulk composition of the sodalite foyaite to approximate that of the magma from which the agpaitic part of the ilímaussaq complex formed. from zr and be data it was estimated that, in order to have progressed from the augite syenite stage to the sodalite foyaite stage, 80 to 95% crystallisation of augite syenite (or benmoreite) magma must have occurred and accordingly engell postulated a very large underlying magma chamber. the pulaskite, foyaite and sodalite foyaite units differ texturally as well as petrographically. the pulaskite is coarse-grained and essentially homogeneous, consisting of alkali feldspar, fayalite, hedenbergite, titanomagnetite and apatite with minor nepheline (larsen 1976). in con7676 trast, the foyaite (c. 20 m thick) is very heterogeneous and exhibits both modal and textural layering (fig. 95). layers of pegmatite, c. 1 m thick, occur at the tops of the layers, with crystals that have grown perpendicularly downwards. each pegmatite then grades into normal coarse foyaite below (ferguson 1964; larsen & sørensen 1987). the underlying sodalite foyaite marks the onset of agpaite crystallisation. because the composition of the sodalite foyaite is similar to that of the calculated average agpaite it may approximate to the magma composition (ussing 1912; sørensen 1958, 1969; gerasimovsky & kuznetsova 1967; engell 1973). the sodalite foyaite is coarse but more evenly grained than the units above, with a poikilitic texture in its lower parts. the rock comprises alkali feldspar, nepheline, alkali pyroxene, alkali amphibole and sodalite, with minor early-formed hedenbergite, fayalite, titanomagnetite and apatite, and late analcime and natrolite; in addition the characteristic agpaitic phases eudialyte and rinkite are present (ussing 1912; ferguson 1964; hamilton 1964; larsen 1976). according to hamilton (1964), the modal percentage of fayalite decreases downwards through the sodalite foyaite. as the pulaskite-foyaite-sodalite foyaite succession accreted downwards, mg in the melt decreased whilst na and zr concentrations increased (larsen 1976). the temperature is estimated to have fallen from c. 900 to c. 800°c. initially the magmas were in equilibrium with a h2o-free high-temperature mineral assemblage (alkali feldspar, fayalite, hedenbergite, ti-magnetite and apatite). the change from foyaite to sodalite foyaite involved an increase in nepheline as well as the appearance of sodalite. the latter signalled the stage at which the magma became saturated in chloride. when the temperature fell to c. 700°c, volatile saturation is thought to have been attained and exsolution of a fluid phase brought about reaction of the high temperature mafic minerals to alkali amphibole, aegirine, aenigmatite and eudialyte (larsen & sørensen 1987). naujaite. the sodalite foyaite is underlain by the remarkable rock type which ussing (1912) called naujaite. although the modal assemblage of the naujaite is essentially identical to that of the foregoing sodalite foyaite, the texture and mineral proportions are strikingly different. naujaite, unique in composition and texture, contains sodalite as the dominant component. the sodalite is typically present up 40–50% modally but can vary from 20 to 75% (sørensen 2006; fig. 96). the crystals (2–3 mm large) appear in two distinct morphologies (hamilton 1964; larsen & sørensen 1987), dodecahedra and hexagonal prisms. the latter, formerly thought to pseudomorph nepheline (hamilton 1964), are more probably paramorphs after a high-pressure polymorph (a.a. finch, personal communication, 2012). ussing (1912) recognised that the concentration of idiomorphic to euhedral sodalite must have been brought about through a flotation process, a conclusion accepted by all subsequent investigators. the density of sodalite (c. 2.29 g/cm3) is presumed to have been less than that of its host magma so that the sodalite primocrysts floated up and accumulated beneath the sodalite foyaite. fayalite, hedenbergite, titanomagnetite and apatite are fig. 95. layered foyaite overlying naujaite, between tunulliarfik and kangerluarsuk fjords. 77 still present as early phases (primocrysts), but are only very minor components in the naujaite. being dense phases the bulk of them may have sunk to contribute to an unseen complementary floor sequence of cumulates (larsen 1976). the naujaites are very loosely compacted cumulates with the intercumulus taking the place of the 30–60% contemporary melt. most of the latter crystallised to alkali feldspar, nepheline, arfvedsonite and eudialyte oikocrysts up to 30 cm across (fig. 96). their large size is taken as further evidence that the melt had very low viscosity, providing exceptional ease for ionic migration. the feldspar is mainly microcline microperthite (with some cryptoperthite) marginally altered to analcime and natrolite. the thickness of the naujaite unit is estimated at some 600 m (andersen et al. 1981a) but, as its lower levels have been magmatically eroded by later magma, the original thickness is inferred to have been significantly greater (sørensen 2006; sørensen et al. 2006). on the assumption that the naujaite extended across the whole agpaite complex, a volume of >60 km3 has been estimated (sørensen 2006). a large, convecting, slowly cooling magma chamber with crystallisation along its roof, walls and floor is envisaged (larsen & sørensen 1987). sodalite crystallising at depth along the chamber walls may have ascended to contribute to the downgrowing roof cumulate. furthermore, in order to account for the extraordinary quantity of sodalite in the naujaites, a parental magma chamber with a volume many times greater than that of the present volume of the complex is supposed (larsen & sørensen 1987; rose-hansen & sørensen 2002). such a chamber must have had a volume at least ten times greater than the estimated (minimum) 60 km3 of the naujaite. large-scale layering in the naujaite (fig. 97) is due to the occurrence of conformable pegmatite horizons about 0.5 m thick (ussing 1912; larsen & sørensen 1987), separated by 10–30 m of normal rock. in the pegmatite horizons the sodalite crystallised downwards from the contemporary roof as prismatic crystals. this, together with the pegmatites in the foyaite, affords a second example of inward-growing crescumulates, the development of which may have coincided with periods of tranquillity when convection in the underlying magma diminished, allowing volatile concentration beneath the chamber roof. addition of volatiles is presumed to have lowered the melt density to less than that of sodalite, temporarily preventing further flotation of sodalite primocrysts, i.e. there were interludes when the normal process of accretion ceased and sodalite crescumulates developed in situ. thus, the naujaite crystallisation front appears to have accreted downwards in a pulsatory fashion. there is a generalised increase in the amount of pegmatite down-section, signifying concentration of volatiles in the diminishing host magma (larsen & sørensen 1987). another form of layering in the naujaites noted by hamilton (1964) is that in places mafic layers composed of arfvedsonite and aegirine “are not uncommon”. thin layers (c. 25 cm) of mafic rock containing concentrates of prismatic aegirine pass upwards into normal naujaite in which aegirine has poikilitic morphology. there are also unusual features in the naujaites of narsaq elv (northfig. 96. naujaite outcrop showing one or more giant oikocrysts of eudialyte (reddish brown). hammer shaft c. 50 cm long. 7878 west ilimmaasaq, fig. 90) in which masses of naujaite (up to 2 m) are enclosed in another naujaite facies (hamilton 1964). clearly there are some features within the naujaites suggestive of more complex marginal structures. larsen & sørensen (1987) mention discontinuous screens of naujaite within the younger pegmatite zone surrounding the kakortokites that may represent remnants of a former marginal facies to the naujaite body. geochemical changes in the upper roof sequence include marked differences in the zr/u ratios through the pulaskite, foyaite, sodalite foyaite and higher parts of the naujaite, and a much lower concentration of u for any given zr value in the lower naujaites (bailey et al. 2001). clearly naujaite growth was not a simple steady-state process but one of considerable complexity. in these sodaliterich rocks, chlorine is a major rock-forming element; the naujaites typically contain 2–3.5 wt% cl but the content can reach 4.6 wt% (bailey et al 2001). cryptic layering, previously noted in the pulaskite-foyaite-sodalite foyaite suite, persisted in the naujaites. early sodalites are enriched in br, i and b relative to later ones (bailey 2006). sulphur is also present in the sodalite as so4 (the sulphatic sodalite referred to as hackmanite) although both sulphide and sulphate ions coexisted in the melt (krumrei et al. 2007). the cores of the sodalite crystals contain minute aegirine prisms as well as hydrocarbon inclusions; study of the latter reveals that the host sodalites grew from a highly reduced, halogen-rich magma in equilibrium with ch4 at c. 800°c. the sodalites are inferred to have acquired their aegirine and hydrocarbon inclusions in the course of their crystallisation during magma ascent. by contrast, their inclusion-free rims may represent crystallisation during emplacement (krumrei et al. 2007). whereas the naujaites crystallised at a pressure of c. 1 kb, the fluid inclusions in the sodalites are deduced to have been trapped at pressures of up to 4 kb (krumrei et al. 2007). on this basis, the depth of the magma in which the sodalites commenced growth could have been as much as 12 km. accordingly, the crystals may have grown over a wide range of depths, either during passive ascent (flotation) or while they were entrained in rising magma. this conclusion necessitates a re-assessment of the hitherto accepted belief that the agpaitic rocks at ilímaussaq have a total thickness barely exceeding 1 km (sass et al. 1972). the conduit through which the naufig. 97. terracing in naujaite on the southern side of tunulliarik fjord, caused by in-weathering of pegmatitic horizons. 79 jaite magma ascended may have been restricted in size (dyke-like?), spreading laterally into a near-horizontal tabular body (c. 100 km2, sørensen 2006) at or near the eriksfjord formation basal unconformity, as is inferred for the giant dyke intrusions. such a geometry may explain the discrepancy between a model demanding a very large volume of agpaitic magma crystallising at depth and the conclusion reached by sass et al. (1972). however, the heat-flow measurements by sass et al. (1972) were made at kvanefjeld near the north-west margin of the complex; this opens the possibility that, had the measurements been taken at a more central locality, significantly higher values might have been obtained, more compatible with the concept of a much larger (deeper) phonolitic chamber capable of supplying the great quantity of sodalite requisite for naujaite formation. it would be of future interest to investigate fluid inclusions within the sodalites grown in situ in the pegmatitic layers in the naujaite. floor series accumulated on a hypothetical floor beneath the naujaites are the kakortokites that occur in an excellently exposed succession that must rank among the most astounding examples of layered cumulates on the planet. the kakortokites compose the lowest exposed 300 m of the succession and pass gradationally up into c. 400 m of lujavrites. the stratigraphy was established by bohse et al. (1971) and reviewed by bohse & andersen (1981) and andersen et al. 1981; fig. 98). kakortokites. the kakortokites are generally separated from their wall rocks (augite syenite, julianehåb granite and eriksfjord formation quartzites) by a steep pegmatitic zone, 25 to 100 m wide (andersen et al. 1981a; sørensen 2006) that is absent from most of the western, northern and eastern agpaite contacts and is essentially restricted to the lower part of the complex, adjacent to the kakortokite and aegirine lujavrite. the pegmatite zone is texturally heterogeneous, with fine-grained fig. 98. layered kakortokites at kringlerne, looking south across kangerluarsuk fjord. the unlayered mass in the centre, immediately above a talus slope, is a large xenolith of naujaite. in the far distance are peaks of the julianehåb batholith (redekammen) behind the southern margin of the complex. 8080 foyaite intervening between abundant anastomosing veins of pegmatite. the former may provide insight into the nature of the magma from which the layered series crystallised (sørensen et al. 2006; sørensen 2006). the outer boundary of the marginal pegmatitic zone is sharp but the inner side (against the kakortokites) is more indistinct (bohse & andersen 1981). it would appear that a relatively finer-grained chilled facies was extensively modified by later migration of volatiles, down a temperature gradient, towards the chamber walls. analogy may be drawn with the marginal border group of the eastern syenite at the kûngnât complex (upton et al. 2013) and comparison may also be made to the outflow of residual fluids from the igdlerfigssalik syenites (see below). the marginal zone surrounds twenty-nine well-defined layered units, dipping gently (c. 10°) towards the centre of the intrusion and making up the lower part of the exposed sequence. these units, composing the lower layered kakortokite series (fig. 98), are numbered from −11 upwards to +17 and have thicknesses of 3.5–12.5 m with an average of c. 8 m (bohse et al. 1971). the idealised unit is tripartite, commencing abruptly with an arfvedsonite-rich base of black kakortokite that grades up into increasingly eudialyte-rich red kakortokite succeeded in turn by white kakortokite in which feldspar and nepheline are the dominant components. however, in some units the red kakortokite facies is poorly developed or even absent. the normal grading in each unit (fig. 99) has been accepted by most investigators as explicable in terms of gravitational sorting during crystal settling, reflecting the decreasing density in the sequence arfvedsonite, eudialyte, feldspar+nepheline. the kakortokites are orthocumulates in which the principal cumulus components are alkali feldspar, nepheline, eudialyte and arfvedsonite. fluorite and aenigmatite attain cumulus status in some units (l.m. larsen 1977; sørensen & larsen 1987). thus the kakortokite magma appears to have crystallised along a remarkable poly-component cotectic. arfvedsonite as a cumulus phase is, however, restricted to the (black) basal layers and the mineral is only present as an intercumulus component in the red and white layers. there is a significant difference in the degree of compaction from bottom to top of the units, with close-packing of arfvedsonite in the basal portions grading to loose packing in the unlaminated white kakortokite tops (upton 1961). from their thin tabular morphology the microcline microperthitic feldspars of the kakortokites inferentially crystallised as monoclinic sanidines. the transition from white tops of the units to the overlying black bases can take place over several centimetres or can be knife-sharp (ferguson 1964). apart from localised thinnings (as beneath roof rock autoliths), the units tend to retain constant thickness and to be laterally continuous for distances of c. 5 km along strike. although there are some indications of incipient trough erosion and deposition in the lower units there is generally very little evidence for convective flow of the magma. the passage upwards in each unit from well-laminated mafic or ultramafic bases to unlaminated leucocratic tops, attributed to progressively declining flow of magma currents by upton (1961), more probably reflects close-packing of the dense arfvedsonites contrasting with low degree of packing in the felsic tops. the observations suggest that the kakortokite magma was remarkably tranquil and had low viscosity, thus permitting virtually complete settling of all arfvedsonite crystals as each unit commenced crystallisation. the agpaitic magmas are estimated to have been de-polymerised and at least as fluid as basaltic magmas despite their lower temperatures (larsen & sørensen 1987; bailey et al. 2001). accordingly they would have been capable of turbulent flow so that, initially, only a small percentage of the primocrysts (cumulus) could remain in suspension. because eudialyte was part of the cumulus assemblage, zr was a compatible element and the zr content of the melt may have been consistently reduced from a maximum of c. 9000 ppm (bailey et al. 2001). the kakortokites (and the succeeding lujavrites) lack the fayalite, hedenbergite, titanomagnetite and apatite of the roof series, implying that the magmas from which the kakortokites and lujavrites crystallised were more fig. 99. graded units in kakortokite of the ilímaussaq complex. scale 50 cm. 81 evolved than those that yielded the roof series (larsen 1976; bailey et al. 2001; sørensen & larsen 1987). by the time the lowest exposed kakortokite was crystallising, some 800 m of the roof series had already formed (bohse & andersen 1981). although the amphiboles in the kakortokites are more magnesian and less calcic than those of the naujaites (larsen 1976) and the floor series growth was separated from that of the roof series by a significant time lapse (steenfelt & bohse 1975; sørensen et al. 2006) this does not necessarily imply that two distinct intrusions were involved, merely that the temporal equivalents of the roof series in the floor series is well below unit −11. the magma chamber is considered to have been tabular with an estimated length and breadth of 17×8 km (andersen et al. 1981a) and a relatively shallow depth variously estimated at >1 km (bohse & andersen 1981) to <1 km (pfaff et al. 2008). although visible only in the relatively uplifted southern part of the complex, the assumption is that the kakortokites extend right across the whole agpaitic complex. the low dips of the units steepen close to the margins so that overall, the layering defines a wide basin-like geometry with upturned margins with dips up to 50°. bohse & andersen (1981) suggested that this form reflects an original sedimentary feature (fig. 100). from studies of other gardar intrusions (cited above and in upton et al. 1996) this appears highly likely. whereas the principal heat loss is assumed to have been through the roof (larsen & sørensen 1987; krumrei et al. 2006), some heat loss through steep sidewalls promoting crystallisation would have led to foot-wall cumulus deposition and accretion of inward-dipping crystal talus. this does not deny that some late-stage floor sagging may have contributed to the geometry (ussing 1912; bohse & andersen 1981). traced laterally, the kakortokite units grade into the marginal pegmatite, their regular black, red and white layers become thinned, broken and folded as they merge into the matrix to the pegmatite zone. within this matrix cross-bedding, graded bedding and wash-out channels are recorded (sørensen 2006), with the implication that dynamic action was sufficient for magma flow to erode previously deposited cumulus and winnow the minerals. large autoliths of naujaite, up to several hundred metres across, accompanied by inclusions of augite syenite and foyaite, occur at one main horizon (unit +3) and are regarded as resulting from a major roof collapse (ferguson 1964; bohse et al. 1971; sørensen 1978). they demonstrate that naujaite was already at or below its soliarfvedsonite lujavrite lujavrite transition zone aegirine lujavrite ii aegirine lujavrite i transitional layered kakortokite slightly layered kakortokite lower layered kakortokite ka ko rt ok ite s lu jav rit es marginal pegmatite augite lujavrite julianehåb batholith fig. 100. cross-section through the margin of the ilímaussaq intrusion, showing steepening of kakortokite and lulavrite units against the marginal pegmatite. inclusions of naujaite and augite syenite within the layered sequence are indicated diagrammatically. modified from bohse & andersen (1981). 8282 dus by the time the lower layered kakortokites were accumulating (bohse & andersen 1981). whilst there are no impact structures beneath the autoliths, they were capable of compressing the underlying unit to about half its thickness (ferguson 1964). the implication is that the unit under pressure was incompletely crystallised and that loading caused expulsion of c. 50% modal intercumulus melt. this is disputed by bailey & gwozdz (1994) who consider that the intercumulus melt content was as low as 15%. the tabular, slab-like form of the autoliths conforms to a pattern common for inclusions in a number of the gardar plutons (e.g. klokken and western kûngnât complexes). they tend to form conformable horizons suggesting detachment along low-angled joints in the roof rocks, thus permitting magma to ascend to a higher level. emplacement of the kakortokite-lujavite magma may have followed such a pattern. despite the low density of its sodalite component, the solidified naujaite must have had a density greater than that of the underlying kakortokite and lujavrite magmas for the autoliths to sink. there is a broad analogy with the situation at syenitknold (see fig. 48) where a roofing facies apparently generated by plagioclase flotation had, after reaching its solidus, acquired a whole-rock density greater than the evolved (trachytic) magma in the underlying chamber, permitting it to sink after breaking off. at a later stage in the magmatic evolution of ilímaussaq the density of the iron-rich lujavrite magma is deduced to have increased to match that of the naujaite so that autoliths failed to sink but remained more or less static in the luvavrite. autoliths of naujaite within the succession as a whole tend to increase in abundance upwards (bohse & andersen 1981; fig. 100). slump structures in kakortokite units −6, −5 and −4 are probably products of gravitational sliding of unconsolidated mafic cumulus from steep sidewalls (fig. 101). the slumps indicate the depth of unconsolidated cumulus to have been at least 20 m (bohse & andersen 1981). relatively steep to very steep sidewall dips are seen in several other gardar intrusions, e.g. the ygdc described above and the kûngnât complex (upton et al. 2013). slumped cumulates are known from the ygdc and also from the nunarssuit syenites (upton et al. 1996). by analogy the kakortokite slumps may also have originated from gravitational instabilities in steep sidewall cumulates. origin of the kakortokite layering. there have been numerous attempts to explain the layering in the kakortokites (e.g. ussing 1912; ferguson & pulvertaft 1963; ferguson 1964, 1970a). one explanation, proposing repeated convective overturn of the magma (bohse et al. 1971), was dismissed on the grounds that the thin tabular geometry of the chamber would not have supported such convection. subsequently an elegant model was proposed involving double-diffusive convection and the upward crystallisation of a compositionally stratified magma (larsen & sørensen 1987; sørensen & larsen 1987). the model appeals to differences in nucleation and growth rates between different minerals. the mafic minerals nucleate and grow at lower degrees of undercooling than the felsic ones. the sharp unit boundaries may correspond to sudden increases of volatile pressure and/or increase in degree of undercooling triggering crystallisation in the overlying magma layer. possibly the separation of large amounts of feldspar and nepheline led to increase in the vapour pressure, releasing heat that, in turn, increased the degree of undercooling. thus crystallisation of successive stagnant bottom layers resulted from differing degrees of undercooling of the minerals in a multiply saturated magma (sørensen & larsen 1987; larsen & sørensen 1987). this successive layer by layer crystallisation in response to upward loss of heat (and some volatiles) was responsible for the overall uniformity of mineralogy and chemistry. pfaff et al. (2008), however, considered that magma layering induced through double diffusion would have yielded only thin (centimetre-scale) layering. problems arise in explaining why the vapour pressure was increased and how the model accounts for the marginal steepening of the layers. the model also encounters difficulties in explaining the bowl-shaped disposition fig. 101. slump structures in kakortokite close to the southern shore of kangerluarsuk. hammer c. 40 cm long. 83 of the layering and evidence for some magmatic flow as well as the gently undulatory form of the units and their draping over the naujaite autoliths. another unanswered question is to what extent were the magma layers crystalline when they were generated. were they aphyric or bearing only microcrysts? a contrasted mechanism suggested by pfaff et al. (2008) was based on the concept of geyser eruptions, namely that a gas phase separating from a magma in a closed system increased its hydrostatic pressure. when the latter exceeded the lithostatic pressure, volatiles were released and, as the vapour pressure fell, the lithostatic pressure promptly sealed the vent and closed the system. as crystallisation recommenced, vapour pressure increased and the process was repeated many times. this model is a refinement of earlier ideas on vapour pressure control to explain the repetitive layering in the kakortokites (ussing 1912; ferguson & pulvertaft 1963). the most serious objection to these hypotheses is that volatile retention appears to have been complete up to the final stage of crystallisation of the agpaite magma. pfaff et al. (2008) proposed repeated influx (multiple replenishments) of new magma after each vapour release event, suggesting that the lower layered kakortokites crystallised not from a single overlying magma body but from numerous batches supplied from a large underlying chamber. an oscillation between closed and open system conditions is envisaged. in order to have produced the c. 8 m kakortokite units, it was deemed necessary to postulate a magma body c. 600 m deep. lindhuber (2011), however, noted that, although at the base of unit +7 of bohse et al. (1971) there is evidence for flow across an incompletely solidified surface (of unit +6), it is only at this horizon that there is clear evidence for the influx of new (slightly more primitive) magma. because of this lack of evidence for replenishment at the bases of the other units, lindhuber (2011) invoked the concept of ‘mineral crowding’ in which rapidly sinking arfvedsonite crystals catch up with smaller (slower) crystals beneath, generating amphibole-rich mats, a process that may have occurred simultaneously at different levels, forming distinct physical barriers. the amphiboles and eudialytes exhibit a sympathetic cyclicity of compositional change upwards through the stratigraphy. for the arfvedsonites, this is seen only in the black basal layers, the only part of each unit in which this mineral was cumulus. fe2+/ mn in the eudialytes decreases from the bottom to top of each unit, indicating progressive fractionation in the magma in each unit. with the exception of units 0 and +7, the ratio is essentially constant in the black layers but progressively decreases up through a unit’s red and white layers (lindhuber 2011). however, the lateral continuity of the units over long distances is difficult to reconcile with the crystal mat concept. as not infrequently observed in science, an initially simple hypothesis (ussing 1912) has been shown by subsequent research to be erroneous: as more data accrue, the more complex the phenomena appear. despite now being in possession of far more detail on the field relationships, chronology, mineralogy and geochemistry, a satisfactory explanation for this fascinating cumulate succession that lacks significant cryptic layering and has such striking macro-rhythmic layering, is still awaited. kakortokite-lujavrite transition. the lower layered kakortokites are overlain by approximately 50 m of poorly exposed kakortokites. although layering in these is indistinct, prominent modal layering reappears in the overlying c. 60 m of transitional layered kakortokites. the highest layered unit of these has an aegirinerather than arfvedsonite-dominated base but, as this is overlain by ‘red’ and ‘white’ layers as in the lower layered series, it is still regarded as kakortokite. these transitional kakortokites pass conformably up into lujavrite cumulates. whereas in the lower layered kakortokites compositional changes in whole-rocks and minerals (specifically arfvedsonite and eudialyte) are small, such changes become much more pronounced in the overlying strata. this phenomenon had been noted with respect to the upward increase in u in the eudialytes (bohse et al. 1974; steenfelt & bohse 1975) and zr/y ratios (andersen et al. 1981b). more recent work has demonstrated an upward decrease in ca/(na+k) in the arfvedsonites and in ca/ (ree+y) in the eudialytes; these changes are much more accentuated in the lujavrites than in the kakortokites. there is a marked decrease in fetot/mn and in the range of compatible trace elements in the rocks, whilst the incompatible trace elements increase (pfaff et al. 2008). these phenomena reflect strong fractional crystallisation in a diminishing volume of magma at the latest stages of ilímaussaq evolution. chlorine, br and i contents had become exhausted by persistent sodalite crystallisation in the first agpaite event but the f content reached a maximum in the kakortokite stage before decreasing during the lujavrite stages (bailey et al. 2001). there is a continuum from the kakortokites up through the overlying series to the highest lujavrites beneath their naujaite roof (rose-hansen & sørensen 2002 and references therein). lujavrites. the lujavrites are defined as melanocratic, eudialyte-bearing nepheline syenites and, whereas the 8484 dominant ferromagnesian mineral in the kakortokites is arfvedsonite, aegirine is commonly dominant in the lujavrites. despite having major minerals in common, the kakortokite and lujavrite suites differ in their minor mineral components and in grain-size and texture, the lujavrites being finer-grained and more fissile (ferguson 1970c). the lower part of the lujavrite sequence is aegirine-rich whereas arfvedsonite predominates in the upper parts (bohse & andersen 1981; sørensen et al. 2006). the lujavrites differ from the kakortokites in that both albite and microcline co-existed as discrete phases (ussing 1912), implying a change from hypersolvus conditions during kakortokite deposition to subsolvus conditions for the lujavrites, brought about by falling temperatures and rising vapour pressure, i.e. there was a significant change in the physical conditions of crystallisation. the thickness of the lujavrite sequence has been variously estimated from 300 m to >500 m (sørensen 2006) and the measured thickness in the southern part of the complex is 485 m (andersen et al. 1981a). the rocks possess a steep lamination close to the contacts but the lamination is approximately horizontal in the more central parts of the intrusion (bohse & andersen 1981). although mainly trapped between the kakortokite and naujaite, the lujavrite transgresses its eastern and western contacts north of tunulliarfik to intrude both the roof series and the eriksfjord formation lavas. aegirine lujavrites in the lower part of the sequence retain cumulate textures but these are less well-developed than in the lower layered kakortokites. still younger lujavrites appear to represent very poorly compacted cumulates, the bulk composition of some being thought to converge with that of the melt. the principal components are aegirine, nepheline, microcline, albite, eudialyte and analcime so, as described above, transition from kakortokites to lujavrites involved change from hypersolvus to subsolvus crystallisation. the aegirine lujavrites have been subdivided into a lower group, ‘aegirine lujavrite i’, with grain-size decreasing upwards, and an upper group, ‘aegirine lujavrite ii’, characterised by large arfvedsonite oikocrysts (bohse & andersen 1981). this is finer grained and lamination and fissility are less extreme than in aegirine lujavrite i. the faint layering in aegirine lujavrite i is absent in aegirine lujavrite ii (bohse & andersen 1981). a time gap between aegirine lujavrites i and ii was marked by intrusion of quartz syenitic sheets that cut aegirine lujavrite but which are themselves cut by arfvedsonite lujavrite (rose-hansen & sørensen 2001). this observation is of interest in demonstrating that at this late stage in the rift system, some silica oversaturated magma was still available. a transitional succession (60 m thick) separates aegirine lujavrite ii from the overlying main arfvedsonite lujavrite layer (150 m thick). above the aegirine lujavrite ii is a 20 m thick lujavrite unit characterised by centimetre-sized ‘augen’ of nepheline or eudialyte. still higher in the succession, alternating layers of aegirine lujavrite and arfvedsonite lujavrite are common (rose-hansen & sørensen 2002; fig. 102). sodalite, nepheline, albite, microcline, eudialyte and aegirine are present as cumulus phases in the transitional lujavrites, with arfvedsonite generally confined to the intercumulus. according to ferguson (1964, 1970c), some layers in both the black and green lujavrites display density stratification. in some places the green/black layers could be due to infiltration of arfvedsonite lujavrite along planes in the aegirine lujavrite, but elsewhere the transformation from one to the other was either gradational or episodic (rose-hansen & sørensen 2002). the crystallisation of either aegirine or arfvedsonite was determined by the activities of water, silica and fo2 (larsen 1976; markl et al. 2001), and the alternation between arfvedsoniteand aegirine-rich layers in the arfvedsonite lujavrites may at least partly be related to pressure relief caused by fracturing of the naujaite roof (rose-hansen & sørensen 2002). micro-rhythmic layering in the lujavrites draped around naujaite autoliths (fig. 103) commonly involves dark layers up to 15 cm thick alternating with thinner, lighter coloured layers. the dark layers are isomodal. the fig. 102. alternating layers of dark arfvedsoniteand greenish grey aegirine-lujavrites. south-eastern ilímaussaq complex. hammer c. 50 cm long. 85 lower parts of the lighter layers are enriched in nepheline and eudialyte whereas the upper parts are richer in analcime and ree phosphate minerals (bailey et al. 2006). spheroidal bodies up to 20 cm diameter occur locally in the arfvedsonite lujavrites. their sharp, meniscus-like margins suggest that they originated through liquid immiscibility. bulk compositions of the spheroids and their host are similar although h2o contents are lower in the former. the spheroids have distinct rims and cores: the rims contain analcime, brown aegirine and k-feldspar (?adularia) whilst the cores are mainly of arfvedsonite and analcime. the internal differentiation of the spheroids is attributed to very late-stage migration of h2o and k-rich fluids from the interiors to the rims (sørensen et al. 2003). there is little stratigraphic variation of whole-rock zr/u and zr/y ratios within the kakortokites, but these ratios illustrate a marked cryptic variation within the lujavrites (andersen et al. 1981). in the lujavrites u rises steadily in relation to zr and there is a lesser, but complimentary, behaviour of y (fig. 104). overall the zr/u ratios decrease upwards from 1200 in the early black kakortokites to 9.2 in the final naujakasite lujavrites on kvanefjeld. in detail, there are separate zr-u trends in the lower layered (black) kakortokites and transitional layered kakortokites, whilst in aegirine lujavrite i the trends define well-defined stratigraphic intervals. overall, zr/y ratios decrease from 18.2 in the black kakortokites to 2.8 in naujakasite lujavrite, and again there are distinct trends in individual rock types and stratigraphic intervals. each trend is controlled by fixed contents of u and y in their main host mineral, the cumulus (na-zrbearing) eudialyte. the shifts in chemistry are attributed to the onset of crystallisation in a sequence of progressively less dense, liquid layers in the magma chamber (andersen et al. 1981b; bailey 1995; bailey et al. 2001). it is notable that the shifts in magma chemistry within aegirine lujavrite i do not coincide with the macro-rhythmic units in this rock type (bailey 1995). it has long been accepted that the lujavrites represent a normal continuation of differentiation beyond the kakortokite stage (e.g. bohse et al. 1971). two recharge fig. 103. autolithic slabs of pale naujaite enveloped by dark grey lujavrite. note ductile deformation of foliated lujavrite between the slabs. north shore of tunulliarfik. 8686 10 zr/u zr/y 100 1000 1051 50 0 m m-c lujavrite lujavrite arfvedsonite black white red aegirine lujavrite ii aegirine lujavrite ii aegirine lujavrite i transition zone transitional layered kakortokite slightly layered kakortokite lower layered kakortokite al tit ud e ab ov e lo w es t e xp os ed le ve l 800 700 600 500 400 300 200 100 arfvedsonite lujavrite ferrorichterite arfvedsonite, leakeite ferroacitinoliteferroedenite hastingdite, pargasite 2 1 0 6 7 8 olivine atoms per formula unit pyroxene amphibole atoms per formula unit aegirine diopside hed. hedenbergite mg si ca nyböite augite syenite pulaskite and foyaite sodalite foyaite naujaite kakortokite lujavrite katophorite fe/20 ca + mn fig. 104. zr/u and zr/y (whole-rock data) vs. stratigraphic height in the southern ilímaussaq complex. from andersen et al. (1981). fig. 105. compositions of mafic minerals from ilímaussaq augite syenite and agpaites. from marks & markl (in press). 87 events, however, took place within the lujavrite sequence. the arfvedsonite lujavrites in the southern part of ilímaussaq form a sill-like complex with a feeder zone at their base (ratschbacher et al. 2011). andersen et al. (1981a) and rose-hansen & sørensen (2002) concluded that the lujavrites did not consolidate in one continuous chamber but rather in several shallow subchambers that are probably connected with each other. data collated on the mafic mineral chemistry of the augite syenites and agpaites (marks & markl in press) help to illuminate the question of separate magmatic influxes. figure 105 shows the compositional trends of the olivines, pyroxenes and amphiboles. whilst there is no question of the separateness of the augite syenite and agpaitic intrusions, the general conformity of the olivine and pyroxene trends harmonises with the conclusion that these magmas were closely related, although there is clearly a mismatch in their respective amphibole trends. in fig. 106 the generalised petrogenetic coherence of the agpaites is brought out by the composition of eudialytes from the floor sequence in southern ilímaussaq (marks & markl in press). the fe/mn, ree+y, zr/hf and cl contents plotted against height may be taken as a crude reflection of the evolution of the post-naujaite magmas, e.g. in demonstrating their overall increase in ree and decrease in cl. the up-section decrease in cl is caused by the compatibility of this element within the sodalite and eudialyte structures. the greater incompatibility of mn2+ relative to fe2+ in the cumulus assemblages is also emphasised. zr/hf values stayed rather constant until the crystallisation of the late lujavritic residues when the ratio increased. the discrepancies in these four sets of data in the lower part of the arfvedsonite lujavrite succession (arfv-la) might be ascribed to magma recharge from a deeper reservoir. hyperagpaites the most extreme compositions, designated hyperagpaites, occur in the roof of the lujavritic body, particularly in the kvanefjeld area in the north-western part of the complex. a steady stream of residual melts and volatiles towards the top is envisaged by rose-hansen & sørensen (2002). these authors report that dykes and sheets of the uppermost lujavrites penetrate the naujaite. intermittent emplacement of the lujavrite melts took place in vertical zones during several periods of deformation and faulting of the roof (augite syenite and lavas), producing an intrusion breccia (sørensen et al. 1969, 1974; nielsen arfv-l b aeg l llb arfv-l a aeg-l lla aeg-ll tlk (slk) llk fe/mn (molar) ree+y (atoms per formula unit) zr/hf (by weight) ? ? cl (atoms per formula unit) 00 16 12 8 4 0 -4 -8 3 6 9 12 15 0-5 1 1.5 2 0.5 21 1.5120 0 0 20 40 60 80 100 300 250 200 150 100 50 0 -50 -100 -150 -150 -250 proposed fault st ra tig ra ph ic po sit io n (m ) fig. 106. stratigraphic plot of fe/mn, ree+y, zr/hf and cl compositions of eudialytes from the ‘floor series’ agpaites of the southern ilímaussaq complex. for the lower layered kakortokites (llk) only (cumulus) eudialytes from the black layers are represented. as yet there are no data from the slightly layered kakortokites (slk). note: marks & markl (in press) postulate a fault separating the slk from overlying transitional layered kakortokites (tlk). data from pfaff et al. (2008), lindhuber (2011) and ratschbacher (2011). from marks & markl (in press). 8888 & steenfelt 1979). the lujavrites in the kvanefjeld area may represent an offshoot from the larger lujavrite body in the central area. crystallisation of the kvanefjeld lujavrites led to an explosive release of volatiles and forceful intrusion of volatile-rich magma into fractures where it crystallised as hyperagpaitic rocks such as naujakasite lujavrite (sørensen et al. 2011). the hyperagpaites possess a wealth of na-rich minerals and, in the latest sodium-rich residua, nepheline became unstable and was eventually replaced by naujakasite, na6(fe,mn)al4si8o26, which can form up to 75% (modal) of the rocks (khomyakov et al. 2001; sørensen & larsen 2001; andersen & sørensen 2005; sørensen et al. 2011). villiaumite (naf), in place of fluorite, is confined to the highest levels in the lujavrites and is a characteristic mineral of the hyperagpaites. a further reaction involved growth of steenstrupine, na14(ce,th,u)6 mn2fe2zr(po4)7si12o36(oh)2·3h2o), at the expense of eudialyte. uranium and th ions that had formerly been preferentially accepted by eudialyte-series minerals were now accommodated in steenstrupine, denoting an increase in the stability of th-u complex ions. the behaviour of u indicates that other processes such as formation of u complexes and volatile transfer probably operated (bailey et al. 1981a; rose-hansen & sørensen 2002). thus the hyperagpaites are petrographically unique rocks comprising albite, naujakasite, steen strupine and villiaumite (fig. 107). hidden layered series the lowest unit accessible in the kakortokites is identified as unit −11 (bohse et al. 1971). without drill cores it is necessary to speculate on the nature of the underlying rocks. as noted above, the kakortokites and lujavrites crystallised from more evolved magmas than those of the roof series. the reasonable assumption is that an upgrowing layered cumulate suite grew concomitantly with the downgrowing roof series. accordingly, strata in the hidden layered series contemporary with the sodalite foyaite and naujaite should contain a cumulus assemblage of alkali feldspar, nepheline, fayalite, hedenbergite, titanomagnetite and apatite (larsen & sørensen 1987; sørensen et al. 2006). sodalite, also co-crystallising, may have been selectively lost by flotation to contribute to the downgrowing flotation cumulate. because of the great volume of sodalite encapsulated within the >500 m thick naujaite, the agpaitic magma body must, as emphasised earlier, have been many times greater. consequently the corresponding floor cumulates should be several kilometres thick. beneath these a downward progression of floor cumulates should be expected, equating to the pre-agpaitic foyaite and pulaskite stages. pursuing this hypothesis, the foyaite and pulaskite cumulates should logically be underlain by a layered series, accumulated at >700°c, of augite syenite and thence syenogabbro as seen in the ygdc cumulate sequence. it is tempting to equate the hidden layered series to the upgrown succession in the older giant dyke (ogdc on tuttutooq) which shows the sequence augite syenite – pulaskite – foyaite – sodalite foyaite. although the ogdc is a much smaller and finer-grained intrusion than its ilímaussaq counterpart, the affinity is obvious. not inconceivably, a small-scale sodalite cumulate may be present in the hidden upper, peralkaline facies of the ogdc, beneath the waters of narsaq sund. in the naujaites, feldspar, nepheline, olivine, titanomagnetite and hedenbergite crystallised in equilibrium with the early sodalites and these minerals may well have been contributing to contemporary cumulates in the hypothesised hidden series of upgrowing cumulates on the floor of the early agpaitic magma chamber. it is noteworthy that, much earlier in the older gardar, a magmatic progression through augite syenite, pulaskite, foyaite to agpaite had come about in the motzfeldt complex (jones & larsen 1985). fig. 107. photomicrograph of hyperagpaitic lujavrite at kvanefjeld. albite (colourless), naujakasite (greenish-grey), steenstrupine (opaque) and villiaumite (orange). 89 ilímaussaq parental magma as described earlier, the ogdc contains a succession grading up from augite syenite through pulaskite and foyaite to peralkaline sodalite foyaite. although its uppermost facies must lie offshore and is unavailable for study, the sequence can be regarded as the approximate inverse of that seen in the downgrown roof of the ilímaussaq agpaites. as the ogdc salic magmas were regarded as residual from fractional crystallisation of the basalt magma that initiated the younger gardar southern rift system, it is deduced that the ilímaussaq magmas were also products of a related basaltic parent magma (upton et al. 1985). the ilímaussaq whole-rock compositions demonstrate a hundred-fold increase in u and th from the initial augite syenite to the final lujavrite stages, requiring >99% crystallisation of the starting magma. furthermore, a continuous fractionation process within a chamber below the present outcrop level is implied because the th-u data show no gap between the augite syenite and the earliest nepheline syenite (pulaskite) (bailey et al. 1981a). the whole complex is considered to be ultimately derived from a single basaltic parental magma fractionating in the deep crust but with the silica-oversaturated intrusive phase requiring crustal assimilation (larsen & sørensen 1987; stevenson et al. 1997; marks et al. 2004). ilímaussaq εnd values range from −0.9 to −1.8 and oxygen δ18osmow is 5.2 to 5.7‰. these data are taken to indicate derivation from an isotopically homogeneous oib-type mantle source, generally without indication of crustal contamination except for the alkali granite which has a lower εnd of −3.1, probably due to contamination in the lower crust. the mantle source is inferred to have been slightly enriched in 18o but depleted in 147sm (marks et al. 2004). the oxygen fugacities of the parental melt were below the fmq buffer curve, resulting in highly reduced mineral assemblages (karup møller 1978; konnerup-madsen et al. 1979). two immiscible fluids separated and were present through most of the crystallisation. one was a ch4-dominated vapour and the other a highly saline aqueous fluid (petersilie & sørensen 1970; konnerup-madsen 2001; krumrei et al. 2007). however, as mentioned above, the claim that the methane (and other hydrocarbons) are of mantle origin has been disputed by laier & nytoft (1995, 2012). micro-kakortokite dyke a dyke immediately to the south of the ilímaussaq complex provides unique insight into the nature of the magma(s) from which the agpaites grew. this ene– wsw-trending, 10–30 m wide dyke, traceable for c. 18 km, transgresses the ilímaussaq augite syenite at its southern extremity (fig. 108; larsen & steenfelt 1974). although it has been noted in earlier sections that the late gardar sinistral faulting coincided with the end of extensional rifting, this, together with the late basic dykes, shows that the cessation was not absolute. the dyke consists of porphyritic phonolite with tabular alkali feldspars, an0.0ab53.4or46.6, (up to 7 cm large, and composing c. 10% (modal)), accompanied by nepheline, fig. 108. map showing the relationship between the southern part of the ilímaussaq complex and the micro-kakortokite dyke (thick dashed line trending ene–wsw). modified from allaart (1969) and larsen & steenfelt (1974). re de ka m m en 10 00 m julianehåb fjord 3 km ilímaussaq complex eriksfjord formation julianehåb batholith dolerite dyke microsyenite dyke micro-kakortokite dyke qaqortup im aa tunulliarfik n ka nge rlu ars uk 60°55´ 45°50´ 9090 hedenbergite and fayalite (fo0.8fa92.7te5.1la1.4) and magnetite microphenocrysts. although it is considered that the dyke magma was initially homogeneous the dyke rocks exhibit two contrasting facies; viz. high alkali and low alkali. in the first, the matrix comprises microcline, albite, nepheline, sodalite and natrolite together with aegirine, arfvedsonite (grown around fayalite), aenigmatite (at the expense of magnetite), eudialyte and fluorite. consequently, the petrography links it indelibly to that of the sodalite foyaites of the roof series, with which it was probably coeval. the affinity to ilímaussaq is confirmed by its wholerock composition which is very much the same as that of average kakortokite. marks & markl (2003) generally concur with these conclusions, noting that the microkakortokite magma separated at an early stage from the ilímaussaq agpaite magma chamber. the whole-rock content of alkalis (na2o+k2o) is 13–15 wt% in the high alkali facies but 11–13 wt% in the low alkali facies. the latter is petrographically distinct, e.g. in being devoid of eudialyte (the zr being accommodated in hiortdahlite and zircon) and is accordingly classified as miaskitic rather than agpaitic. the difference between the two facies is ascribed to alkali loss during emplacement and crystallisation (larsen & steenfelt 1974). the localised loss of na2o and h2o as well as of f, cl and some trace elements into the wall rocks in lowtemperature fluids can be considered in the same light as that from comendite dykes (tugtutôq), carbonatite dykes (with fenitised margins) in the nunatak region and, on a larger scale, the fluid loss from the igdlerfigssalik syenites. the micro-kakortokite dyke magma had low fo2 and it was concluded that, in order for undersaturated salic melts to generate a characteristic agpaitic mineral assemblage, they had to be iron-rich, strongly peralkaline and capable of retaining their alkalis (larsen & steenfelt 1974). østfjordsdal syenite and igdlerfigssalik complex age relationships the timing of the østfjordsdal syenite and the igdlerfigssalik complex presents a dilemma relative to the tugtutôq and ilímaussaq complexes. the østfjordsdal syenite on the south-east side of the igdlerfigssalik complex is clearly older than the latter (fig. 67). both the østfjordsdal syenite and the younger components of the igdlerfigssalik complex are intersected by members of the ene-trending dyke swarm. the whole of the igdlerfigssalik complex appears to have been emplaced before movements along the left-lateral c. e–w transcurrent faults ceased (emeleus & harry 1970). although this prima facie evidence suggests a greater age for these two relative to the tugtutôq and ilímaussaq complexes, this is contradicted by the radiometric ages (table 1) that indicate rb-sr ages of 1148 ± 3.6 and 1142 ± 15 ma for østfjordsdal and the late igdlerfigssalik complex respectively. the østfjordsdal rb-sr dating is supported by a u-pb (zircon) age of 1147.5 ± 3.2 ma (table 1). thus on the basis of the age determinations these two could be c. 10 ma younger than ilímaussaq and therefore among the youngest intrusions in the province, together with the recently discovered paatusoq intrusion (1144.1 ± 1.1 ma, table 1) situated well outside the rift zone on the southern contact of the julianehåb batholith, some 90 km east-south-east from igdlerfigssalik. østfjordsdal syenite although truncated by one of the latest igdlerfigssalik units, the østfjordsdal syenite appears to have been a subcylindrical stock with a diameter of c. 5 km. it is largely composed of coarse-grained syenite consisting of alkali feldspar, nepheline and subordinate aegirine-augite and biotite. it is also cut by some trachyte and lamprophyre dykes whilst being younger than some phonolitic dykes (emeleus & harry 1970). so far very little has been published on the østfjordsdal syenite. igdlerfigssalik complex the igdlerfigssalik complex, like its south qôroq predecessor, has an elliptical plan (11 × 15 km), elongate south-east–north-west and, as at the south qôroq complex, the faulting may have been responsible for this geometry, deforming the rocks while they were still hot and ductile (stephenson 1976b). it is the southernmost major intrusive centre amongst the igaliko syenites. the complex cross-cuts the south qôroq syenites as well as the østfjordsdal syenite that lies on its extreme south-east side (fig. 67); emeleus & harry 1970). 91 igdlerfigssalik, which reaches a height of 1752 m, is well-exposed with deep dissection (fig. 109). the physical difficulties, however, imposed by the terrain are such as to leave much of this complex unstudied in detail (emeleus & harry 1970). the seven syenite intrusions that compose most of the complex are divisible into two groups, three older and four younger. using the (modified) symbols for the intrusions as divided by emeleus & harry (1970), the first of these groups comprises i1, i2 and i3 which pre-dated the latest stages of the ene dyke intrusion. only a narrow strip of i1 remains, in the north-western part of the complex, and i2 is also fairly vestigial, seen as a strip on the northern side of the complex. the outcrop of i3 around the north side is less than 1 km wide but has a well-developed inward dipping lamination at 20–70°. the majority of each of these units has been obliterated by its successor. the second group of syenites comprises i4, i5, i6 and i7 and was emplaced after the intrusion of the igaliko dyke swarm had come to an end, implying that the igdlerfigssalik complex was not complete until after the regional lithospheric extension ceased. eriksfjord formation strata south-west of i4 (in the tunulliarfik and igaliku fjord areas) are approximately horizontal up to a c. 1 km broad zone adjacent to the intrusion, in which the strata are flexed downwards at up to 60° towards the contact. this may reflect downsag (drag) towards the syenite that accompanied subsidence of an approximately cylindrical block of the country rocks as i4 was being emplaced. intrusion i4 crops out around the north, west and south sides of the complex, displaying contrasting facies. it is inferred to have been part of a steep-sided stock involving a marginal border group and an inner layered cumulate series subsequently largely replaced by i5, i6 and i7 (emeleus & harry 1970). amongst the several facies of i4, the ‘dark layered syenite’ is considered to have been part of a marginal border group. as implicit in its name, this shows prominent development of modal layering brought about by concentration of ferromagnesian minerals. the layering is fig. 109. view east towards nepheline syenite mountain of the igdlerfigssalik complex. in the foreground and middle distance julianehåb batholith overlain by outliers of the eriksfjord formation, visible beyond the farm. see also frontispiece. 9292 steeply inclined with layers cut (eroded) by layers farther from the outer contact, providing good evidence for inward younging (fig. 110). these rocks were compared by emeleus & harry (1970) to those of the eastern border group at the kûngnât complex (upton et al. 2013) where the layering is ascribed to magma downflow alongside thermal boundary layers. well-laminated, concordant cumulates in i4 are relicts of the former layered series. this contains xenolithic masses of anorthositic gabbro. these, mantled by overlying syenite cumulates, generated impact disturbances in the underlying syenites. to find such rocks in the salic gardar intrusives is unusual, the great majority being in mafic and intermediate hosts. as the protoliths are believed to be of lower or mid-crustal origin, their presence may indicate rapid uplift in an earlier magma batch before they collapsed into the upgrowing i4 cumulates. i4 is cut by sheets of syenitic and microsyenitic rocks which themselves contain xenoliths of both the i4 host syenite and anorthosite or gabbroic anorthosite. a pegmatite close to narsarsuk on the southern border of i4 and related to a porphyritic microsyenite sheet has been a mecca for mineralogists. one mineral among the assemblage, named from the locality, is narsarsukite, na2(ti,fe)si4(o,f)11 (flink 1901). intrusion i5 forms a broad annular outcrop (fig. 67) accounting for most of the complex including the summit. it is a remarkably uniform, coarse-grained syenite but, where signs of layering appear, all dip towards the intrusion centre. i6 forms a virtually complete 360° ring dyke (c. 35 km circumference) mostly separating i4 from i5. it is ovoid, 15 × 10 km diameter with a long axis trending nw–se, varying in width from 600 m to 30 m. the rock textures are highly variable from mediumgrained to coarse-grained with pegmatitic patches. i7, the youngest of the syenitic units, is ovoid in plan (7 × 5 km) and lies entirely within i5. it is a steep-sided stock with contacts dipping outwards at 75–80°, consisting of a medium-grained leucocratic foyaite. structurally it consists of thick layers stacked one on top of the other in a shallow saucer-like form with gently upturned margins. it presents a superb section across a layered intrusion, displaying conformable igneous lamination and small-scale fig. 110. inward-dipping layers in syenite unit i4 of the igdlerfigssalik complex exposed in the cliffs above qooroq fjord. 93 modal layering (fig. 111). thus i5 and i7 form a central downfaulted block bounded by the i6 ring-dyke. cumulus phases in the igdlerfigssalik foyaites are alkali feldspar, clinopyroxene, apatite, magnetite ± olivine and nepheline; intercumulus phases include zonal overgrowths to the above, plus amphibole, biotite, aenigmatite, sodalite, nepheline, alkali feldspar, alkali pyroxene and magnetite. additionally subsolidus phases were produced through the action of h2oand/or co2-rich fluids acting on the magmatically crystallised minerals. subsolidus products include biotite (grown through alkali feldspar–magnetite–h2o interaction), blue-green amphiboles around olivines, (silica and alkali rich relative to the browner intercumulus amphiboles), cancrinite through interaction of co2-rich fluid and nepheline, as well as gieseckite from nepheline and sericite from feldspar (powell 1978). whilst the cumulus assemblages in the foyaitic rocks of the south qôroq and igdlerfigssalik complexes are more evolved and complex than those in the ygdc and the klokken complex, very similar processes operated within the magma chambers. the rocks are dominantly poorly compacted orthocumulates reflecting the relatively rapid cooling of magmas at shallow depths. olivine–clinopyroxene thermometry indicates crystallisation between 980° and 900°c, consistent with watersaturated liquidus temperatures at 1 kb (powell 1978). the fluted sidewall cumulates in igdlerfigssalik unit i4 point towards flowing magma currents with the capacity to thermally or mechanically erode previously formed layers (fig. 112). in the igdlerfigssalik magma chambers, as in those of several other gardar plutons, it is probable that convection involved two-phase downflow driven by loading of high-density pyroxenes ± olivines. the infig. 111. photo and sketch of layered structures in syenite unit i7 of the igdlerfigssalik complex as seen from the south-eastern side of qoororsuaq, looking west towards the complex. sketch modified from emeleus & harry (1970). 1505 m i7 i7 i7 qoororsuaq gravel flats scree southern lake 1550 m 570 m i5 9494 ward inclination of the modal layering, most perfectly preserved in unit i7 of igdlerfigssalik, can again be assigned to the accumulation of cumulus minerals deposited peripherally as crystal talus or pediment by crystalrich slurries detaching from the thermal boundary layers before flowing radially towards the central part of the chamber floors. the high degree of parallelism widely observed in the tabular feldspar crystals of the laminated syenites is ascribed to orientation by flowing magma as it was in the ygdc gabbros. whereas in the south qôroq complex there was progressive evolution towards successively more primitive magma batches (stephenson 1976a), there are, as yet, no data to discern any such pattern for the igdlerfigssalik intrusions. late fluids expelled from the late igdlerfigssalik intrusions migrated outwards to form an aureole c. 1 km wide. this transects the south qôroq complex and is marked by a zone in which the south qôroq ferromagnesian minerals are recrystallised (stephenson 1976a). finch (1995) demonstrated that the fluids emanating from the igdlerfigssalik complex reacted with the biotites of neighbouring rocks, specifically affecting the octahedral sites and hydroxyl sites. the unaffected south qôroq rocks contain biotite with a fluorine content reflecting that of the late-stage south qôroq fluids. by contrast, in the recrystallised aureole, the f content of the biotite is distinct, inferentially closely related to that of the late-stage fluids expelled from igdlerfigssalik. these fluids also modified the ree, zr and hf contents of the affected south qôroq biotites. exchange of late-stage fluids, exuded from younger intrusions, may be commonplace in alkaline igneous systems and the f content in the micas provides a sensitive indicator of this phenomenon (finch 1995). eccentric to all of these salic components is a partial ring dyke of gabbro/syenogabbro that crosscuts the eastern side of the complex and transgresses the i5, i6 and i7 syenites as well as the østfjordsdal syenite (fig. 67). consequently this can be compared to the very late intrusion of mafic magma (from a source below that of the syenitic magmas?) that occurred in the south qôroq complex. because igdlerfigssalik rises to >1.7 km above sea level and its lower outcrops cut the eriksfjord formation that was probably never more than 4 km thick, the uppermost parts of the complex clearly penetrated high in the eriksfjord formation. it would therefore be surprising if the igdlerfigssalik complex did not have an extrusive expression. i6 may, for instance, have broached surface level to produce an eruptive curtain around the subsiding block of i5 and i7. the exposed syenites may have been components within a composite body that lay either within higher stratigraphic levels of the eriksfjord formation or possibly within the superstructure of a large overlying volcano. assuming the forms of the annular syenites approximate to the sizes of calderas in a nested suite, such a volcano (built up largely of pyroclastic products?) may have had a diameter from 50 to 75 km. possible modern analogues include kilimanjaro in tanzania and cantal in auvergne. considered jointly, the south qôroq and igdlerfigssalik complexes may illustrate consecutive attempts to build a large salic volcano over an active tectonic zone. products of the first attempt then experienced more or less concurrent faulting and dyke fissuring before being overgrown by the second (igdlerfigssalik) volcano that was affected in its early phases by dyke intrusion as lithospheric extension reached its close, and which probably post-dated all but the very latest stages of the transcurrent faulting (stephenson 1976b). the role of anorthosite anorthosites of the nain province are major features in the mesoproterozoic geology of labrador, but do not outcrop east of the labrador sea. nonetheless, the abundance of anorthositic xenoliths in gardar intrusions leaves no doubt that an extensive anorthosite body underlies the province, constituting a petrological ‘elephant in the room’: obvious but rarely discussed. it is inferred to have accreted over the whole time-scale of gardar magmatism and to have played a seminal role in their petrogenesis (bridgwater 1967; bridgwater & harry 1968; upton 1996, halama et al. 2002). fig. 112. cut-and-fill cross-bedding in the igdlerfigssalik syenite unit i4. scale 1 m long. 95 the distribution of xenoliths suggests a parent body with estimated dimensions of 250–500 km by 50–100 km, i.e. comparable in size to the angola and nain (labrador) anorthosites (emslie 1977). both the cryptic gardar anorthosite and the mid-proterozoic anorthosites of labrador and quebec lie close to a terrane boundary and were related to extensional tectonics and failed rifts (morse 1982, 2006). there are numerous parallels between the michikamau anorthosite of the nain province and observed or inferred features for the sub-gardar anorthosite. at michikamau the succession entails anorthosite, ferrodiorite, ferromonzonite and ferroadamellite (emslie 1965, 1970) whereas the deduced gardar succession is anorthosite, ferro-syenogabbro, ferro-syenite, syenite (and thence foyaite or alkali granite). the principal differences between the nain and gardar provinces appear to be the much shallower level of erosion and the more alkalic characteristics of the latter. the gardar rocks overlie or intrude the julianehåb batholith, but there are relatively few data regarding the make-up of the archaean–palaeoproterozoic lithosphere beneath the gardar province (see the discussion in garde et al. 2002 and references therein). seismic studies offshore south greenland indicate a moho depth of 30–35 km. dahl-jensen et al. (1998) interpreted seismic reflection data to suggest a thick wedge of archaean continental crust under the eastern part of the batholith, whereas a pb-pb isotopic study in the west by kalsbeek & taylor (1985) showed that an archaean lead isotopic signature at the north-western margin of the batholith quickly disappears towards the centre of the batholith. garde et al. (2002) concluded that the bulk of the batholith (and hence also the deep crust of presumed mafic composition) consists of juvenile material that was accreted onto the southern margin of the archaean craton. this leaves the depth of the presumed anorthositic gardar cumulate open to interpretation. xenoliths of anorthosite and gabbroic anorthosite, together with plagioclase megacrysts, are especially abundant within the mafic and intermediate intrusions of the two younger gardar rift zones (bridgwater & harry 1968). halama et al. (2002) conducted studies on isotope and trace element geochemistry of the megacrysts in the isortoq district, confirming that most of the anorthosite xenoliths are alkaline and cognate with the gardar magmatism. the maximum pressures deduced for these megacrysts are 10–12 kb. others, however, may have crystallised higher in the crust, suggesting a polygenetic and polybaric history. in comparison, the nain complex in labrador may consist of over 20 different plutons (wiebe 1992). anorthosite formation may have taken place throughout most of the gardar time. the gardar anorthosite body is presumed to predate the giant dykes and the main dyke swarm. bridgwater (1967) and bridgwater & harry (1968) suggested that granular anorthosites (with a specific gravity of c. 2.63 g/cm3 at 1000°c), formed as flotation cumulates deep in the crust and acted as a trap for residual magmas (fig. 75). the feldspars are extensively sericitised; this, and secondary oxidation of the oxides, was attributed to accumulation of water and other volatiles close to the anorthosite before entrainment. the granular anorthosites were thus thought to have formed a more or less impermeable cap above the differentiating alkaline magmas. generation of so large a body, with its restricted composition, must have involved repeated batches of relatively evolved magma. given the sodic labradoritite composition and the overall alkaline nature of the xenoliths, the parental magmas would have had to be hawaiitic. although plagioclase cumulates may have formed temporary chamber roofs that had some mechanical strength, it is probable that such low-density bodies at depth were unstable and occasionally were disrupted during tectonic disturbances, yielding crystal-liquid mushes. these would subsequently crystallise to coherent rocks capable of fragmentation and entrainment in basaltic to intermediate magmas. most of the granular xenoliths approximate to pure anorthosite of remarkably constant composition, with plagioclase of an61-56 (bridgwater & harry 1968). the feldspar crystals are typically equant, anhedral and randomly oriented. crystal sizes are mainly from 3 to 5 cm but can exceed 20 cm. in order of decreasing modal importance, other phases are olivine (fo73-60), fe-ti oxides and pyroxene, typically augite. blocks of texturally distinct granular anorthosite within host granular anorthosites points to a complex polycyclic origin for the protolith, with possible additional mixing from diapirism at depth. the earlier components may show some deformation, which is attributed to compaction rather than to tectonic deformation. laminated anorthosites the laminated anorthosite autoliths are considered to be genetically distinct from the granular anorthosites 9696 (bridgwater 1967; bridgwater & harry 1968). they appear to be confined to the southern rift, principally in the ygdc and klokken gabbros, and consequently are considered to have played an important role in the magma genesis of the system as a whole. the laminated anorthosites have a less complex history than the granular anorthosites and retain pristine and perfectly preserved cumulate textures (fig. 58). whilst there is superficial similarity between the granular and laminated anorthosites, their mode of formation was clearly contrasted. while the former probably accreted as flotation cumulates (see above), the latter are more likely to have grown as cumulates upwards from a magma chamber floor. by analogy with the nain and angola examples, the laminated type at asorutit may represent a layered body that overlies granular anorthosites. the problem of accounting for both supposed floor and roof cumulates in the nain anorthosite complex was called the ‘feldspar/ magma density paradox’ (morse 1973; scoates 2000). the laminated xenoliths in the ygdc at asorutit have idiomorphic tabular plagioclase (an58-52) enclosing intercumulus olivine (fo71-69), subordinate augite, ilmenite, biotite and apatite. the relatively small haematite component in the ilmenites (upton & thomas 1980) indicates a reduced state of oxidation whilst the lack of strong zonation in the plagioclases and their coarse grain-size relative to that of the troctolite host suggest that cooling was slow in comparison to that of the host magma. there is wide disparity in sizes of the plagioclase crystals, suggesting that nucleation occurred at different levels in the parent magma. the latter is inferred to have approximated to a plagioclase-olivine-melt composition. some of the xenoliths (up to 100 m across) show primary layering in which well-laminated layers alternate with poorly compacted layers with more randomly oriented plagioclases and correspondingly greater amounts of olivine. consequently there is modal layering but it is solely due to textural changes; plagioclase remains the sole cumulus participant. the absence of cogenetic olivine+plagioclase and olivine floor cumulates is attributed to their greater density. whereas the xenoliths with sufficiently high plagioclase content would float, the more mafic cumulates could not. closely similar phenomena have been described from the michikamau anorthosite, labrador (emslie 1970) and the paul island anorthosite, labrador (wiebe 1992). the anorthosites at asorutit were interpreted as having originated as floor cumulates in which the degree of lamination was controlled by the rate of magma flow, good laminar orientation being related to flowage whereas the disorientated layers were attributed to tranquil interludes (upton 1961). the original attitude of the hypothetical chamber floor is, of course, unknowable. scoates (2000), however, considers that sloping floors may be characteristic of all proterozoic anorthosite complexes to allow downslope drainage of relatively high-density intercumulus melts. since virtually all the floor cumulates exposed in the gardar province are inclined, a sloping floor for e.g. the assorutit anorthosites, is an acceptable proposition. genesis of the anorthosites among basic igneous rocks, it is uncommon to find evidence for olivine crystallising late relative to plagioclase and, on the basis of behaviour in the an-fo binary system (osborn & tait 1952), upton (1961) proposed that a fall in ph2o had shifted the cotectic towards olivine, thus extending the plagioclase field. consequently, plagioclase nucleated alone until fractionation brought the system back to the olivine-plagioclase-liquid cotectic when olivine precipitation commenced. the relative compositions of plagioclase and olivine differ significantly from troctolite host to anorthosite xenoliths, i.e. from an65 and fo68 in the former to an58 and fo71 in the latter, possibly attributable to pressure differences pertaining at the different crystallisation depths. furthermore, the whole-rock incompatible element ratios (assumed to reflect those of the melts for these orthocumulates) also differ. thus p/ti and p/zr ratios of the anorthosites are notably higher than for the host troctolites. high whole-rock contents of sr (1700–1800 ppm) suggest distinctly high sr concentrations (and relatively low ca/sr) in the anorthosite parent magma. chondritenormalised ree patterns (fig. 113) are generally similar to those of the initial ygdc and ogdc magmas but are more strongly fractionated, with la/ybn c. 18.1 in the laminated anorthosites versus c. 11.5 for the ygdc and c. 16.9 for the ogdc. it is of interest that the normalised patterns for the two asorutit samples closely resemble those for proterozoic orthopyroxene-bearing anorthosites (wiebe 1992). from the plagioclase composition a hawaiitic parent magma is presumed, with high concentrations of lree, p and sr. whilst the ygdc, ogdc and inferred anorthosite magmas were all cognate, they represented quite distinct magma batches. basaltic melts are more compressible than crystals at >6 kb, and plagioclase more sodic than an90 then has a density less than melt (kushiro & fuji 1977; kushi97 ro 1980). these authors concluded that, for polybaric crystallisation models, plagioclase may sink at low pressures but float at the higher pressures appropriate to the deeper crust, and the density difference will increase as fe-contents rise from hawaiitic to mugearitic compositions. bridgwater (1967) postulated that the gardar anorthosites formed in the lower or middle crust. experimental studies on ygdc intitial compositions (upton & thomas 1980) showed that the olivine-plagioclasemelt equilibrium did not persist above c. 6 kb, suggesting a maximum depth limit for anorthosite formation of approximately 20 km. study of the laramie (wyoming) anorthosites led mitchell et al. (1995) to conclude that fractionation of magmas in the upper mantle produced high-al basic residues that gave rise to plagioclase-rich diapirs that ascended and crystallised at mid-crustal depths. however, data for plagioclase megacrysts from the isortoq bfds implied crystallisation near the crustmantle boundary (halama et al. 2002). scoates (2000) pointed out that there is little direct evidence for plagioclase flotation cumulates either in layered intrusions or in proterozoic anorthosite complexes. however, evidence favouring flotation of plagioclase megacrysts and anorthosite xenoliths is apparent from the relationships described above from the tugtutôq, narsaq and klokken complexes. whereas the assorutit occurrences are all of laminated anorthosite, both granular and laminated anorthosite inclusions occur in the klokken complex. it has been widely accepted that the proterozoic anorthosites derive from partial melting of depleted upper mantle sources and that the melts pond at or near the crust/mantle boundary (ashwal 1993). more recently a contrasted genetic hypothesis has gained support which proposes that the anorthosites came from reaction between underplating basalts and aluminous lower crust (duchesne et al. 1999; bédard 2001; charlier et al. 2010). the y-yb-lu data for the anorthositic nain plutonic suite indicate residual garnet in garnet granulite source rocks (bédard 2001). light has been cast upon this petrogenetic dilemma by gleissner et al. (2010) from studies on the kunene anorthosite complex of angola. this, one of the world’s largest massif-type anorthosites, comprises two successive intrusions: an older one of pyroxene anorthosite and leuconorite, and a younger one dominated by olivine-bearing anorthosite. the conclusion reached is that the magma of the first was subject to crustal contamination whereas the second, with characteristics close to those inferred for the gardar anorthosite protolith, experienced almost no contamination. the kunene and the inferred gardar anorthosites are similar in size. with respect to kunene, gleissner et al. (2010) concluded that only partial melting of a mantle source could provide enough primary melt to give rise to so large an anorthosite body and that plume activity was probably involved. moreover, extensive melting of the lower crust is unlikely to have been on such a scale as to produce the requisite vast amounts of almost homogeneous parental melts. accordingly, the kunene interpretation supports a simple mantle origin for the gardar anorthosite and there appears no obvious reason to discard the more conventional ashwal (1993) hypothesis. furthermore, the alkaline nature of the gardar anorthosite is not readily compatible with the afc model involving garnet granulite. ro ck / ch on dr ite 100 200 80 60 40 20 10 8 6 4 2 la ce pr nd sm eu gd dy ho er yb lu ogdc ygdc 50221 186217a fig. 113. chondrite-normalised ree patterns for two anorthosite samples from asorutit, compared to ree ranges from the ygdc and ogdc marginal facies (upton 1996). 9898 emplacement mechanisms and tectonics emplacement mechanism of the giant dykes big dykes with widths on the 100–1000 m scale are rarities in the phanerozoic but are comparatively common in the precambrian. assuming that asthenospheric convection was more vigorous in the precambrian, shear rates on overlying lithospheric plates were correspondingly greater. the greater stresses led to more dramatic failures than in phanerozoic rifts and great volumes of basaltic melts, like those of the ygdc, could be concentrated and intruded rapidly as dykes with widths on the kilometre scale (macdonald & upton 1993). the great majority of dykes in the region, both the early gardar bd0 dykes that long pre-date the younger gardar southern rift and also those that post-date the giant dykes, were simple dilational intrusions. the ogdc has essentially parallel margins for its 20 km outcrop. its broad bow-shaped (northwardly convex) plan (fig. 5) presents some problem with respect to intrusive mechanism but this is dwarfed by the much more abrupt ‘tight’ bend, (concave to the north) developed about half-way along its trace. the western termination of the ogdc also merits attention. despite poor outcrop, the western termination of the ogdc appears to be along an e– w-trending plane oblique to the dyke trend, suggesting that the plane had been a pre-existing shear zone in the granitoids that acted as a mini-transform fault during dilation of the dyke fissure. this hypothesis implies that the country rocks on the northern side of the fault were parted by a diagonal displacement of c. 3500 m during intrusion although the width of the dyke was only c. 500 m. the absence of internal chills and typical lack of extensive wall-rock melting in the ygdc imply rapid intrusion and that, if there were any conduits for surface eruptions, these were highly localised. the giant dyke branches are considered to have crystallised as unitary cooling systems. however, the forms displayed on the geological map by the ygdc (figs 5, 10) present a significant problem with regard to its emplacement. field evidence shows that the ygdc branches were capable of expansion and contraction in both horizontal and vertical senses. in places the branches underwent a localised expansion for which the term ‘ballooning’ is appropriate. for example, the northern branch of ygdc on tuttutooq, traced west from narsaq sund (figs 10, 17) shows a gradual constriction to a narrow ‘waist’ c. 400 m broad, with some ballooning on either side of this waist. a more pronounced instance of this behaviour is shown some 2 km west of itillip saqqaa where the northern ygdc branch subdivides (fig. 23). the northernmost of the two sub-branches displays some remarkable features. in contrast to the approximately constant width of 500 m maintained for several kilometres to the east, it locally expands, over a distance of few tens of metres, to a width of c. 700 m. this width is maintained for a short distance (c. 1.25 km) before it narrows abruptly into a westward extension a mere 200 m wide. still farther west this progressively diminishes to a few tens of metres before swelling once again before terminating against an approximately e–w shear zone, very much as exhibited by the ogdc. as with the ogdc terminus, this ending (and also that of a parallel giant dyke branch to its south-east), suggests that the shear zone acted as a mini-transform fault. the ygdc outcrops reappear east of ilimaasaq at kangerlua (fig. 10) reaching their fuller expression on mellemlandet and nunataks north of motzfeldt sø. in the nunatak region between nordtop and geologfjeld, 2 3 4 fig. 114. outcrop forms in giant dykes at isortoq. grey: mafic facies. red-brown: salic facies. extraction at 60°55́ n, 47°30´w from nunarssuit geological map at scale 1:100 000 (pulvertaft 1967), with old spelling of place names. the circled dykes 2–4 refer to bridgwater & coe (1970). 99 two lenticular pods of differentiated rocks occur in the southern dyke branch between marginal sheaths of troctolitic gabbro. at syenitknold (fig. 10), the giant dyke made a remarkable shift of course, as if it had been displaced dextrally some 200 m along a wsw–ene fault. despite this appearance, no fault was detected and it appears that the dyke fissure simply made an abrupt change of course. it is relevant here to consider the behaviour of the giant dykes in the northern (nunarsuit–isortoq) rift in order to gain a better perspective on the giant dykes of the southern rift (fig. 114). of the many gardar dykes in the isortoq region, the giant dykes are the youngest and, on the basis of available age data (bangs havn, table l), they are approximately synchronous with those of the southern rift. five (or more) of the isortoq dykes are comparable to the ygdc in having (a) giant dyke dimensions with widths up to 500 m, (b) internal synformal layering, (c) possession, at least locally, of a composite character with gabbroic or syeno-gabbroic outer sheaths (or border groups) enclosing evolved central facies, typically of syenite, and (d) a remarkable propensity to exhibit pinch and swell morphologies. this last characteristic was emphasised by bridgwater & coe (1970) who considered it incompatible with emplacement through simple dilation, claiming that stoping must have been involved. bridgwater & coe (1970) described four of these intrusions on either side of isortoq fjord, from north to south, as dykes 1, 2, 3 and 4. dykes 2, 3 and 4 only are shown in fig. 114. dyke 2 shows some of the most extreme behaviour, breaking up along its length into a series of rounded pods as if boudinaged. they are the youngest dyke intrusions in a terrane that had already experienced intensive intrusion by earlier doleritic dykes. still more aberrant is the highly localised ballooning of one young isortoq dyke from a width of c. 10 m to c. 900 m, generating the lenticular (pod-shaped) bangs havn intrusion with a gabbroic sheath around a syenitic and granitic core. the map (fig. 115) also shows the sinuous, flamboyant forms of the associated dykes. here, as with the intrusions described by bridgwater & coe (1970), it appears more probable that the magmas intruded highly sheared granitoids that were sufficiently hot to yield in a ductile manner in the extensional regime as mafic magma ascended. the giant dyke shapes in the southern rift are deduced to have resulted similarly from intrusion into a hot granitoid basement but one that was at a lower ambient temperature than in the northern rift, thus resulting in less extreme diapiric behaviour. that pinch and swell features are absent in the ogdc but characterise the ygdc may then be attributed to intrusion of the former when the temperature of the crust had not yet been sufficiently raised. it is of interest that mesoproterozoic giant dykes in sweden exhibit comparable idiosyncrasies to these gardar intrusions. thus the halleförs dyke shows a similar composite nature as well as comparable pinchand-swell features (krokström 1936). the emplacement mechanism for the klokken complex remains unexplained. if, as has been argued above, the initial gabbro at klokken was essentially an integral part of the ygdc intrusive event, how did it acquire its ovoid form? xenoliths of country rock granite gneiss occur within it but, in view of the densities of the latter and the inferred density of the gabbro magma (2.80 ± 0.5, mingard 1990), the granite gneiss could not have been stoped and sunk within the magma. whereas entry of the salic magmas into the axial cores of the composite northern and southern rift giant dykes might have taken place as these intrusions continued to dilate, the geometry of klokken precludes this mechanism. figure 116 is an attempt to explain the tectono-magmatic evolution of the narsaq and ilímaussaq complexes. the five hypothetical maps purport to illustrate the intrusion forms as they may have been below the eriksfjord formation unconformity and the narsaq lopolith. fig. 115. anastomosing dykes and giant dykes north-east of the bangs havn intrusion, nunarsuit–isortoq zone. grey: mafic facies. red-brown: salic facies. extraction at 60°50´n, 47°53´w from nunarssuit geological map, scale 1:100 000 (pulvertaft 1967), with old spelling of place names. 100100 the first (fig. 116a) shows the ygdc as it may have been prior to faulting, with its branches on tuttutooq widening and merging east-north-eastwards. the supposed wnw–ese-trending contacts shown near narsaq are based on the orientation of the contact between ygdc gabbro and julianehåb granite on narsaq island. the same contact orientation is seen in the giant dyke occurring north-west of narsaq across tullerunat killit and also shown on tuttutooq west of sejlfjord (fig. 10). figure 116b envisages the situation following initial leftlateral displacement along the ‘narsaq fault’, whilst fig. 116c illustrates the intrusion of the narsaq complex and its intersection with the fault. the supposition here is that the narsaq complex magmas exploited, as did other gardar magmas, the lithospheric weak point caused by intersection of the fault and the rift axis fissuring. emplacement of the narsaq complex would have involved engulfment of substantial amounts of eriksfjord formation rocks, the narsaq lopolith and giant dyke gabbros plus parts of the julianehåb granitoids. it is supposed that magma emplacement and fault displacements were intimately associated in space and time. figure 116d supposes further fault movement post-dating the narsaq complex that shifted part of the ygdc towards the west, so that its outcrop now lies beneath the waters of bredefjord, immediately north of the ilimmasaq peninsula. figure 116e illustrates the final situation following emplacement of the ilímaussaq complex, the focus of which was sited one or two kilometres further east. in the terminal phases of tectonism, as strain energy was finally dissipated, the ilímaussaq complex acquired its elliptical form and the kinked boundary on its northwestern flank. although these speculative cartoons cannot be correct in detail, they should, in overall principle, approximate the actual tectono-magmatic evolution. tectonics within the southern rift there was a tendency in the gardar province for episodes of intrusion and faulting to alternate (emeleus 1964; watt 1968; upton et al. 2003). demonstration of transcurrent displacements principally along left-lateral approximately e–w faults, and right-lateral along approximately n–s faults, is readily shown in a terrain rich in steep to vertical contacts. the evidence for normal faults is unsurprisingly scarce. a b c d ygdc narssaq complex ilímaussaq complex narsaq e n fig. 116. five hypothetical stages in the tectono-magmatic evolution of the narssaq and ilímaussaq complexes. see text for discussion. 101 normal faulting in a terrane affected by lithospheric extension, such as the younger gardar southern rift zone, normal faults parallel to the rifting would be expected to be abundant, as they are e.g. in the afar district of ethiopia. there are many ene–wsw-trending shear planes within the julianehåb granite, e.g. through tuttutooq. whilst there is usually no evidence for lateral displacement along these, it may be suspected that many did have vertical displacement. the coastlines on either side of the 4 km wide bredefjord north of the tuttutooq archipelago are remarkably straight (figs 1, 2); the glacial erosion that generated the fiord was presumably controlled by shear zones in the basement although geological mapping could prove no displacements. it is suggested that faulting through bredefjord may have marked the northern boundary of the rift. the southern side of bredefjord defines the north coast of the ilimmaasaq peninsula, and along this the eriksfjord formation is seen in its fullest expression, c. 3.4 km thick, whereas on the northern side of bredefjord the outcrop is all below the base of the eriksfjord formation. large-scale vertical displacements are implied. any such faulting was older than the younger gardar but whether it was pre-older gardar or occurred between older and younger gardar times is not known. just as the hypothetical bredefjord faults may represent the northern boundary of the rift, the more or less linear ene–wsw trends of the coasts bounding the skovfjord (fig. 1) south of the tuttutooq archipelago can be construed as marking the traces of normal faults along the southern boundary of the southern rift. the eastnorth-east extension of skovfjord, the tunulliarfik fjord, is regarded as the site of major faulting. downfaulting to the north, in pre-ilímaussaq times, of at least 700 m has been postulated (sørensen 2006). the fault for which the best data are available is the kangerluarsuk-lakseelv fault that transects the ilímaussaq complex, subdividing it into a southern portion containing the floor cumulates and a larger northern portion that exposes higher structural levels (fig. 90). it is a hinge-fault, the throw of which diminishes towards the east-north-east from at least 600 m in kangerluarsuk, in the west-south-west, to near zero at appat on the tunulliarfik coast (sørensen 2006). the northerly downthrow relates to successive movements that were: pre-ilímaussaq and post-aegirine lujavrite and possibly also post-arfvedsonite lujavrite (bohse et al. 1971; sørensen 2006). a minor fault on the northern margin of the ygdc at asorutit is interpreted as a normal fault, as mentioned in the section on anorthosite xenoliths. the gabbro on the southern side is inferred to be downfaulted from a higher structural level in which the xenoliths had been concentrated by flotation. further evidence suggests that an ene–wsw-trending normal fault transects the southern part of the tugtutôq central complex. although the direction of throw is unknown, it is a reminder that some extensional faulting may have persisted until late in younger gardar times. transcurrent faulting figure 80 shows the pattern of transcurrent faults affecting the gardar province. the pattern comprises conjugate suites of ene–wswto e–w-trending left-lateral faults and nnw–sseto nne–ssw-trending rightlateral faults (berthelsen & henriksen 1975; upton et al. 2003). it was, however, the left-lateral faults that were critical in the localisation of the major gardar intrusions. these faults are known to have been intermittently active since pre-gardar times, e.g. the 6 km displacement on the laksenæs fault commenced in the ketilidian and continued into late gardar times (fig. 80; henriksen 1960). lying between 60°30´ and 61°30´ and separated by distances of 20–30 km, these faults segmented the southern rift in a manner comparable to that described for the mesoproterozoic midcontinental rift system of north america (green 1992). three of these fault zones are relevant to this bulletin and will be referred to as the northern, central and southern fault zones. from their effect on the ygdc components, the younger gardar displacements were approximately 8 km on the northern fault zone, 9 km on the central fault zone and 1.4 km on the southern fault zone that roughly bisects tuttutooq, i.e. approximately 18.4 km in total. thus the northern fault zone, which had been influential in the emplacement of the older gardar grønnedal-íka complex (emeleus 1964), also controlled the clustering of both the older and younger gardar members of the igaliko syenites. intersection of the central fault zone and the ygdc appears to have focussed emplacement of the narsaq and ilímaussaq complexes. although no major intrusions occurred in conjunction with the southern fault zone crossing tuttutooq, the fault is of interest in showing that whilst movement on the ygdc was only 1100 m, the ogdc was displaced by 1400 m. the time interval between the two giant dyke intrusions is constrained to about 20 ma (table 1). there is, as noted earlier, a differ102102 ence in their orientation ascribed to a change in the stress field, and a change in their palaeopole positions. consequently, the concept that c. 300 m movement occurred during the time interval between the two intrusions is quite acceptable. some 15 km west of narsaq the central of the three sinistral fault zones under discussion traverses the island of tullerunnat killit (fig. 10). alongside the fault, though not itself sheared, is a section of a composite giant dyke (500 m broad) that may have been a component of the broad and geometrically complex culmination of the ygdc that is suspected to have been present before its engulfment (through subsidence) by the narsaq complex (fig. 116c). as noted above the igaliko plutons tend to be elliptical in plan (long axes trending nw–se) whilst those remote from the fault zones (e.g. klokken and the tugtutôq central complexes) are more circular. stephenson (1976b) accounted for these observations by postulating a simple shear model in which the south qôroq and igdlerfigssalik complexes, while still hot, experienced ductile deformation that modified their supposedly initial circular plan towards that of a simple strain ellipse. the originally circular plan of the south qôroq complex was first deformed by two sets of movement on left-lateral faults while it remained ductile. subsequent left-lateral movements along two further fault planes involved brittle fracture and resulted in the present plan of the centre (fig. 117; stephenson 1976b). noting that similar deformation had also occurred in the older gardar grønnedal-íka complex, stephenson (1976b) speculated that the elliptical plan of the ilímaussaq complex may also have resulted from strain related to the central fault zone, despite the fact that significant movements along it had ceased after intrusion of the narsaq complex and the main dyke swarm. west of south qooroq, between bredefjord and tunulliarfik, two major faults composing the northern fault zone (fig. 80) record a left-lateral shift of 6–6.5 km across a 4 km wide zone, all involving brittle fracturing (emeleus & stephenson 1970). the more northerly of this pair generated a 200 m wide crush zone in which syenites and dykes are crushed and sheared and a downthrow to the north is suspected (emeleus & harry 1970). the more southerly fault, well seen east of south qooroq, has a c. 100 m wide crush zone with a sinistral displacement of c. 1 km of a contact between a syenite unit (s2) and basement granite. it also has a probable vertical throw of more than 400 m. movement(s) on the northern fault may be later than those on its southern neighbour (emeleus & harry 1970). according to these authors, the age of its activity should remain open in view of some evidence that movement occurred after formation of one of the late intrusions (i6) in the igdlerfigssalik complex. of the four faults investigated by emeleus & stephenson (1970) between tunulliarfik and qooroq, and east of qooroq, vertical throws discerned from displacements in the eriksfjord formation strata were downwards towards both the north and south. with regard to the central fault across the narsaq complex, a significant downthrow to the north was inferred by hamilton (1964) and a northerly downthrow has been suggested in this work for the southerly fault across tuttutooq in the vicinity of itillip saqqaa. the southern gardar rift was scarcely affected by the nnw–sse to nne–ssw dextral faults that are widespread across the province (fig. 80). one of these faults (trending n–s) , however, is present on the east side of the igdlerfigssalik complex, with a horizontal movement that displaces the contact between units i2 and i3 by at least 400 m (emeleus & harry 1970). s s ss i i i i a b c d fig. 117. progressive deformation and faulting of the south qôroq complex. i: igdlerfigssalik complex. s: south qôroq complex. modified from stephenson (1976b). 103 evolution of the magmatic system of the younger gardar southern rift parental mafic magmas mafic rocks ranging from basalts and dolerites to troctolitic gabbros were produced across the gardar province from earliest to latest gardar times, i.e. for over 100 ma (table 1). olivine dolerite dykes occur in abundance, and gabbro also participates in several gardar plutons (kûngnât, nunarssuit, klokken, south qôroq and igdlerfigssalik). additionally, much of the eriksfjord formation consists of basaltic lavas. analyses of lavas, dykes and chilled marginal samples from the intrusions from across the province suggest that, irrespective of place and time, the mafic magmas had a common compositional affinity (upton 1969; upton & emeleus 1987; upton et al. 2003). they have relatively evolved compositions with the liquids appearing rarely to have >7 wt% mgo. the analyses can be roughly subdivided into four groups: (1) older gardar dykes (bd0 and early gardar dykes from the far west of the province), (2) eriksfjord formation lavas, (3) dykes and gabbros from the northern (nunarssuit-isortoq) rift zone, and (4) younger gardar dykes and gabbros of the southern rift zone. in each group the mg* number (atomic 100mg/(mg + fe2+)), is <50 and olivine compositions are rarely if ever more magnesian than fo70 .these basaltic compositions are typically poor in the diopside component, leaving them relatively rich in plagioclase and olivine so that they crystallise to troctolitic rocks. this characteristic is manifest in high al2o3/cao ratios in the range 1.75–2.40. all these basaltic compositions are distinctly potassic with average compositions for each group having >0.9 wt% k2o. virtually none are tholeiitic or typical alkali olivine basalts but are transitional olivine basalts plotting close to the ol-pl-cpx plane of critical undersaturation in the normative basalt tetrahedron (yoder & tilley 1962; coombs 1963; upton & thomas 1980). geochemical characteristics of the southern rift mafic magmas the older gardar basaltic lavas and dykes, and lavas and dykes from the younger gardar northern rift, have similar minor and trace elements ratios whereas those of the younger gardar southern rift are markedly different. thus the hree/lree and zr/nb ratios in the southern rift are not only significantly lower than those of the older gardar dykes and eriksfjord formation lavas but are also lower than those of the younger gardar northern rift (upton & emeleus 1987). figure 118 is a ce/y vs. zr/nb plot of data from the four gardar basaltic groups as well as from ultramafic lamprophyres, silicocarbonatites and carbonatites (note that these last include gardar samples of all ages). this diagram involves ratios of two pairs of incompatible elements and in each pair one element (ce and nb respectively) is distinctly more incompatible than the other (y and zr). ce and y are proxies for respectively light and heavy ree. whilst the ratios are insensitive to moderate degrees of low-pressure crystal fractionation involving mineral phases likely to crystallise from basaltic magma, they reflect differences in the degree of mantle melting and/or differences in source composition (hardarson & fitton 1991). the younger gardar mafic dykes of the southern rift (tugtutôq–ilímaussaq swarm in the legend) are clearly distinct from those of the older gardar dykes, the eriksfjord formation lavas and the younger gardar dykes of the northern (nunarsuit–isortoq) rift. all the southern rift data fall within the field of ocean island basalts, whereas this is true for only some of the other three basaltic groups and for some of the lamprophyre-carbonatite association. the southern rift data also fall between the fractional melting curves calculated for depleted garnet and spinel-lherzolite mineralogies (not shown), consistent with derivation of their magmas from a melt column extending across the garnet-spinel transition zone in the mantle (hardarson & fitton 1991). the southern rift basalts also have higher contents of ba and sr than those from elsewhere in the province, irrespective of space and time (upton & emeleus 1987; fig. 119). among the major elements, p2o5/tio2 ratios indicate relative p enrichment of the southern rift magmas (fig. 120). these data emphasise the broad compositional unity of older gardar dykes, the eriksfjord formation lavas and the northern rift magmas on one hand but demonstrate the distinctiveness of the southern rift mafic dykes 104104 on the other hand. the relative enrichment of the latter in lree, nb, p, sr and ba might be viewed as due to a smaller degree of mantle melting than in the other three groups. however, it is not associated with variation of the silica/alkali balance as would be expected. accordingly it is interpreted as reflecting higher concentrations of the most incompatible elements in the mantle source beneath the younger rift zone. as similar concentrations are not seen in the older eriksfjord formation lavas in the same area, it is suggested that the higher concentrations in the younger magmas were due to a focussed metasomatic enrichment above an asthenospheric wedge developed after eruption of the lavas but prior to the initiation of the southern rift. analyses of the most mafic chilled facies rocks from the southern rift zone were presented earlier in table 2. magma evolution in the southern rift zone whilst liquid lines of descent can only be indirectly approximated from the plutonic suites, they can be authenticated through study of the smaller, fine-grained dykes that followed the giant dykes. these smaller dykes are considered to be residues from fractional crystallisation of parental troctolitic magmas. the more primitive gabbros crystallised from melts that had either olivine alone or olivine + plagioclase on their liquidus. however, the plagioclase in these rocks is commonly seen as glomerocrysts with a radiating structure giving rise to ‘snowflake’ gabbros (figs 18, 25). such ‘snowflakes’ are regarded, as noted above, as products of rapid crystallisation from magma oversaturated in plagioclase; further evidence for plagioclase oversaturation comes from occurrences of ‘perpendicular feldspar’ crescumulates as in the ygdc (itillip saqqaa) described earlier (fig. 26). the al and sr contents in the presumed magmas, combined with virtual absence of any negative eu anomalies in the ree patterns (blaxland & upton 1978; upton 1996), support the contention that plagioclase fractionation did not occur until the magmas reached relatively shallow crustal levels (possibly <6 km) and that, for at least part of the ascent, olivine was crystallising alone. comparable textural and geochemical evidence for late and rapid crystallisation of plagioclase from the gardar basaltic magmas is found elsewhere in the province, e.g. in the eqaloqarfia dyke of the isortoq area (pulvertaft 1965), in the older gardar kûngnât complex (upton et al. 2013), as well as in all three of the principal groups of lavas of the eriksfjord formation (poulsen 1964). the supposition is that 15 10 5 0 5 10 2015 zr/nb field of oceanic island basalts ce/y gardar lamprophyres and carbonatites younger gardar dykes (tugtutôq-ilímaussaq swarm) younger gardar dykes (nunarssiut-isortoq swarm) older gardar dykes eriksfjord formation fig. 118. incompatible element ratios in phenocryst-poor basic gardar dykes (4–8 wt% mgo) and eriksfjord formation lavas, shown together with ultramafic lamprophyres and carbonatites. modified from upton et al. (2003). 105 had these same magmas been retained for any length of time to equilibrate in the lower crust, plagioclase crystallising from them would have gone to augment the evolving anorthosite. that this did not occur is presumably because of rapid ascent. the anorthosite xenoliths and associated feldspar megacrysts, which are regarded as broadly cognate with the ygdc, provide contrasting evidence for higher pressure plagioclase crystallisation (halama et al. 2002) and lack any evidence for fast growth from supersaturated melts. since the ygdc magma arose from beneath the anorthosite protolith without attaining the olivine-plagioclase-liquid cotectic until it reached the upper crust, the implication is that it ascended fast, becoming increasingly plagioclase oversaturated until plagioclase nucleation commenced. subsequently, when abundant plagioclase separation commenced, the consequent iron enrichment in the relatively reduced magmas led to generation of ferro-mugearitic melts. prolonged fractional removal of ca-bearing feldspars and later pyroxenes from evolving magmas at depth (inferentially in the lower crust) is considered to have caused the development of peralkalinity in the salic residues in accordance with bowen’s (1928) ‘plagioclase effect’, and also through the agency of ‘the orthoclase effect’ (bailey & schairer 1964) whereby the preferential entry of potassium into the feldspars helped to generate per-sodic (agpaitic) residual magmas. the high al/ca ratios of the basaltic magmas were responsible for the delayed crystallisation of pyroxene. in this respect the gardar magmas have much in common with the parental magmas of the north american proterozoic anorthositic intrusions, e.g. that of the kiglapait complex in labrador (morse 1982, 2006). a principal point of difference between the labrador and gardar parental magmas is the higher k content of the latter. the comparatively low silica activity and high k2o of the gardar magmas precluded crystallisation of low-ca pyroxenes and dictated their evolution via hawaiites, mugearites and benmoreites to trachytes and ultimately to peralkaline rhyolites and phonolites. the magmas also had notably high fluoride contents (upton et al. 2003; köhler et al. 2009). this is considered to have conferred an unusual degree of fluidity (low yield strength) facilitating convection among other things. the unusually coarse-grained nature of the gardar plutonic rocks is attributed to depolymerisation of the magmas by fluoride ions. the f-rich character of the gardar magmas, which are similar in this respect to the andean volcanic rocks, provides another pointer to the mantle source having been affected by subductionrelated metasomatism (köhler et al. 2009). in the foregoing chapters evidence has been adduced for the delicate density balance between solids, whether 1500 1000 500 0 1000 2000 3000 younger gardar (t–i–n rift) ba ppm sr ppm younger gardar (n–i rift) older gardar dykes eriksfjord formation 7 6 5 4 3 2 1 0 1 2 3 4 5 p2o5 wt% tio2 wt% younger gardar (t–i–n rift) younger gardar (n–i rift) older gardar dykes eriksfjord formation fig. 119. ba-sr plot of phenocryst-poor basic gardar dykes (4–8 wt% mgo) and eriksfjord formation lavas. modified from upton & emeleus (1987). fig. 120. p2o5 vs. tio2 in phenocryst-poor basic gardar dykes (4–8 wt% mgo) and eriksfjord formation lavas. modified from upton et al. (2003). 106106 these be discrete crystals (e.g. of sodalite), crystal aggregates (e.g. plagioclase-olivine ‘snowflakes’) and their host melts, dictating whether they sank or floated. in the remarkable bfds (big feldspar dykes), packed with plagioclase-rich rocks and crystals, it may be surmised that melt and solid densities were closely matched. crystalrich slurries are judged to have descended from magma chamber sidewalls. in the case of the anorthosite roof vs. floor cumulates it was suggested that different behaviour shown in rocks of similar composition depended on density changes in the melts according to pressure. in numerous instances, the similarity of layered structures in the intrusions to those of sedimentary sequences also points to a remarkable fluidity of the magmas, whether mafic, intermediate or, as in the case of the agpaites, extreme alkaline differentiates. the apparent ease of separation of crystals from melts allowed highly effective fractional crystallisation and production of extreme lithologies. magmatic differentiation in the lower crust it is inferred that great volumes of gardar mafic magmas were underplated at or near the crust-mantle boundary and were ultimately parental to the alkaline salic plutons. in the case of the kenya rift, to which the southern late gardar rift may have had some resemblance, the volume of magma including underplated material has been estimated as c. 934 000 km3 (latin et al. 1993). the kenya rift has been magmatically active for c. 35 ma whilst the activity in the younger gardar southern rift may have covered c. 40 ma (from c. 1180 to 1140 ma) and the magma volumes involved may have been comparable. in the model offered here, a very large volume of magma was generated above a rising asthenospheric mantle wedge fed by rising plume material (latin et al. 1993). around 1180 ma extensional stress culminated in lithospheric attenuation and partial melting of the metasomatised lithospheric mantle along a zone coinciding with the axis of the julianehåb batholith, leading to extensive underplating by primitive basaltic magma. the latter was relatively ca-poor (hence with high al/ca ratio) and k, sr, ba-rich from their inception. according to herzberg (1995), al/ca of melts decreases with increasing pressure of peridotite melting and equilibration, and the high values of the gardar basalts could signify a relatively low-pressure melting regime. fractional crystallisation of olivine (± pyroxene and spinel?) led to a hawaiitic magma crystallising on an olivine-plagioclase cotectic. at the deep crustal levels plagioclase crystallised and, being less dense than the magma, accreted to form a flotation cumulate while olivine sank to yield dunitic cumulates at the base. this situation is envisaged as having continued intermittently throughout the entire period of gardar magmatism, and repetitive influx of new primitive magma batches must be assumed. the concept for genesis of the salic magmas proposed by bridgwater & harry (1968) and summarised in their cartoon (fig. 75) remains generally valid although in need of modification. the bfds are highly instructive with respect to the petrogenesis in the rift system, providing not only the key linkage between the benmoreitic and the hawaiitic/mugearitic magmas but signalling the importance of compositionally stratified magmas at depth and indicating that these developed beneath an anorthositic lid. when continuing extension resulted in crustal failure, the older giant dyke was intruded. slow crystallisation in the interior of this steep-sided tabular, half-kilometre wide intrusion led to the upward growth of its syenite suite from its residual melts. after several millions of years during which some plate rotation occurred, further buildup of transtensional stress gave rise to a second, still more dramatic crustal failure, in conjunction with a greater degree of melting of the same mantle source. after a significant amount of fractionation, a portion of the basalt magma was emplaced as the younger giant dyke complex. since this magma (like its ogdc predecessor) had all the characteristics of being a residue after extensive fractional crystallisation, it is assumed that these magmas were products of a far greater volume of primitive magma. ascent of the magma disrupted part of the deep crustal anorthosite, entraining large and small masses en route that then accumulated by flotation at the top. rare-earth element data (fig. 113) indicate that the primitive magmas from which at least some of the laminated anorthosites were derived, represent smaller melt fractions of the mantle source than the parental magmas of either the ygdc or ogdc. residual magma retained beneath the anorthosite is envisaged as occupying one or more chambers in the lower crust, elongate parallel to the axis of the southern rift. the approximate dimensions may have been 30–50 km long, c. 15 km broad and perhaps a kilometre or so deep. with slow cooling these magmas then underwent compositional stratification. from the bfd evidence this appears to have comprised a hawaitic/mugearitic lower layer overlain by salic (benmoreitic/trachytic) magma with <2 wt% mgo. production of such stratified magma bodies 107 was probably repetitive. as a consequence of the extraction of plagioclase from these relatively reduced magmas during anorthosite genesis, the residual magmas became increasingly fe-rich and correspondingly dense. this resultant density handicap incurred by iron enrichment is presumably the reason why these magmas rarely reached shallow crustal levels (cf. the gardar ‘daly gap’, watt 1966). (the genesis of ferro-syenogabbros in the ogdc and ygdc is explained as shallow-crustal reflections of what occurred on a greater scale deep in the crust). ygdc magma, with a low density relative to these ferich residual magmas, ascended through the anorthositic capping to reach shallow levels. crustal fissuring permitted selective tapping of the stratified chamber(s) by dyke formation as rift extension continued. this process, generally but not invariably, extracted magmas from the salic top of the chamber(s). attainment of the benmoreite composition appears to have marked an important stage in the rift’s evolution, just as it was for the magmatic evolution in the kenya rift (macdonald 2002). only when the residual magmas became sufficiently iron-poor (total iron as fe2o3 <12 wt%) i.e. benmoreitic, did they attain low enough density to ascend through the crust, independent of dyke fissuring. it was benmoreitic magmas that were the preliminary arrivals in the ilímaussaq, igdlerfigssalik and tugtutôq central complexes. reduction of stress energy with time is suggested by the generalised reduction in dyke widths in the main swarm. the igaliko swarm, introduced during the same tectonic phase as the main swarm, may have resulted from a smaller degree of melting of contrasting mantle sources. as extensional strain energy dissipated, a changed stress regime promoting transcurrent faulting was responsible for dyke formation to become increasingly rare. magma chambers of the central complexes in order to extend the model outlined above to embrace the formation of the principal salic centres, it is hypothesised that at c. 30 km intervals along the rift system’s elongate parent chamber, foci developed where collection of buoyant salic residues was particularly concentrated. from field observations among the gardar alkaline complexes it is surmised that ascent took place by repetitive detachments of slabs of roofing rocks, up to 100 m or so thick, that successively became underlain by lowerdensity magmas. the geometry of the slabs may have been controlled by subhorizontal (thermally induced) jointing. evidence from the grønnedal-íka and kûngnât complexes (emeleus 1964; upton 1960; upton et al. 2013) shows this behaviour where roofing consisted of high-grade gneisses. at nunarssuit, the tugtutôq central complex and also kûngnât, the roofing involved supracrustal mafic volcanic cover, whilst at klokken, syenitknold and ilímaussaq ascent of magma involved displacement of coeval roof sequences (‘upper border groups’). by logical extrapolation from what is seen at current erosion levels to deeper levels, it may be assumed that the same mechanism by displacements between roofing slabs (intact or disintegrated) permitted ascent of the syenitic magmas through the crust. the lithospheric thickness along the southern rift axis cannot be known, but from evidence of modern rifts (e.g. the gregory rift, east africa) it may have been as little as 35 km and crustal thickness correspondingly reduced (macdonald 2003). evidence from the kûngnât and ilímaussaq complexes suggests that their magma chambers were situated at depths of c. 3 km. on this line of argument it may have needed only a limited number of such collapse events to raise the magmas from the lower crustal parent chamber to the shallow crust. in the discussion of the ilímaussaq complex emphasis was placed on the great size of the augite syenite magma chamber required to account for the high concentration of incompatible elements in the agpaitic magmas. such a chamber may well have had the form of an extensive accumulative benmoreitic magma in the upper parts of an elongate, compositionally stratified parent chamber as proposed above. there is a contrast between those major salic complexes where the parent magmas at depth became increasingly evolved with time (the tugtutôq, narsaq and ilímaussaq complexes) and those like south qôroq where successive intrusions had progressively more primitive compositions. the igdlerfigssalik complex ended with a mafic partial ring-dyke but there are as yet no data to show whether the previous six intrusions followed a comparable evolution. the closest analogue to the south qôroq intrusive pattern is that of the older gardar kûngnât complex (upton et al. 2013). the first category suggests two or more admissions from a part of the salic upper layer or the ‘master chamber’ that was evolving through continued assimilation and fractionation (tugtutôq and narsaq) or through fractionation with minimal assimilation (ilímaussaq). in the second category, repeated collapses into a stratified parent chamber may have occurred, culminating in ring-fault descent of the already crystallised (and consequently relatively dense) syenitic 108108 components forcing the underlying hawaiitic component to high levels. the one complex that stands alone, geographically and metaphorically, is klokken. it differs from the igaliko syenites (fig. 2) in being silica-saturated, finishing with an oversaturated diorite. it differs from all the other gardar central complexes in commencing with troctolitic gabbro, with younger units passing progressively from unlaminated syenite to the layered syenites of its centre. as has been documented above, the intrusive and crystallisation patterns at klokken and at the differentiated ygdc pods, at asorutit and syenitknold were so similar as to suggest their possible contemporaneity. however, whereas emplacement in the ygdc may be explicable in terms of initial crustal dilation followed by ingrowth of sidewall cumulates and upgrowth of floor cumulates, the cylindrical geometry of klokken is incompatible with a dilational introduction of the gabbroic magma. in starting with a mafic magma it contrasts starkly with all the other central complexes. how the initial cylindrical pluton, inferentially of mafic magma, was emplaced into granitic-gneiss country rocks remains an enigma. genesis of the silica-oversaturated magmas on approaching the minimum melting point on the alkali feldspar join, the salic residues evolved either to the rhyolitic or the phonolitic minima in the qz-ne-ks system (upton 1974). the generally oversaturated main dyke swarm and the generally undersaturated igaliko dyke swarm can be geochemically distinguished by their zr/nb ratios, those of the igaliko swarm having values <5.2 while dykes of the main swarm have higher values. the small negative nb anomalies in trace element patterns and higher 87sr/86sr values of the main swarm dykes are attributed to greater degrees of crustal assimilation (foland et al. 1993). high-temperature fluids rich in alkalis, volatiles and incompatible trace elements, advancing ahead of rising mantle diapirs or plumes, may have a profound effect on the overlying lithospheric mantle and lower crustal rocks, causing fenitisation and varying degrees of crustal melting (woolley 1987). many continental a-type granites may have been generated in this manner (martin 2006). however, whereas this hypothesis may well apply to the northern (nunarsuit–isortoq) rift, for which a greater degree of crustal heating has been proposed above, it has limited applicability in the southern rift. in the latter, quartz trachyte and comendite dykes occur over much of its length but the quartz syenites and alkali granites are restricted to a 30 km long sector between the tugtutôq and ilímaussaq complexes. dyke propagation probably had a very significant lateral vector, but the magmas of the central intrusions are more likely to have ascended more or less vertically from the regions in which they were generated. the crustal assimilation necessary to generate the silica-oversaturated salic melts through afc processes may have been due to the heat of crystallisation from the deep crustal basaltic bodies affiliated to the ygdc. the greatest effects are seen around a narsaq hot spot that is inferred to denote the principal focus of magma genesis. the tectono-magmatism cartoon (fig. 116) shows how faulting may have distanced the ilímaussaq intrusions from the ygdc, thus minimising crustal assimilation in its genesis. it is remarkable that the southern rift system included both the highly reduced, ultra-sodic hyperagpaites and the extremely oxidised, and potassic mela-aillikites of the narsaq region. these petrologically polar-opposites were proximal in both space and time. silica-undersaturated rocks are absent from the northern (nunarsuit–isortoq) rift zone. this zone experienced no less than five swarms of mafic dykes during its gardar history (harry & pulvertaft 1963), and high resultant geothermal gradients may have been characteristic, particularly during the younger gardar, modifying the mechanical properties of the crust and facilitating crustal assimilation. as outlined above, the anomalous geometries of the giant dykes of this zone may be due to hotter country rocks yielding in a more ductile fashion at the time of giant dyke intrusion. this conclusion conforms with field evidence that the contacts do not show the same degree of chilling as those of its counterparts in the southern rift. crystallisation histories whilst the intrusions along the southern rift zone involved a great number of separate magma batches these still compose broadly coherent lines of liquid descent. whilst the overall crystallisation history of what is argued to be a single magmatic system cannot be deduced from any one part of it, it can be discerned from the collective sources of evidence. these present a remarkably complete petrogenetic narrative from simple crystal–melt equilibria exemplified by the most primitive 109 magmas (represented by the ygdc chilled samples) to the astonishingly complex equilibria in the ilímaussaq agpaites. the most primitive magmas (excluding the aberrant aillikitic magmas) were slightly silica-undersaturated and their fractional crystallisation led via trachyte and phonolite to agpaitic and ultimately hyperagpaitic residuals. this undersaturated trend is regarded as the dominant one within the system as a whole, but it was interrupted by the, geographically more localised, silicasaturated/oversaturated salic magmas discussed above. the ygdc basalt was estimated to have intruded at c. 1140°c (upton 1971; upton & thomas 1980). this involved olivine + liquid, joined at shallow crustal levels by plagioclase. titanomagnetite began precipitation at the maximum iron-enrichment stages, joined at much the same stage by apatite. iron, ti, mn and p contents rose to maxima when mgo was reduced to between 2 and 3.5 wt% and the melt had attained a ferro-mugearitic composition. in tholeiitic magmas fe and ti contents commonly peak during the intermediate stages of differentiation which is held to be a consequence of relatively low degrees of oxidation (fenner 1937). in the alkaline ogdc and ygdc, high fe and ti concentrations were similarly reached at mid-stages of magmatic evolution because the oxidation states were low, somewhat below that of the qfm buffer. uncharacteristically for basaltic magmas, clinopyroxene (salite) only joined the crystallising assemblage at a late stage. the salite evolved to ferrosalite and from this na-enrichment towards aegirineaugite proceeded at differing stages of fe-enrichment according to the oxidation state. in the most extremely reduced case (ilímaussaq) the pyroxenes attained nearly end-member hedenbergite composition before there was any significant intake of na (larsen 1976). because of the high k2o content of the parental basalt (c. 1.4 wt%) the feldspars followed a trend from plagioclase through potassic oligoclase and anorthoclase to sanidine. sanidine was accompanied by nepheline in the phonolites with subsequent appearance of sodalite and natrolite at lower temperatures. the feldspars crystallised under hypersolvus conditions until, with rising ph2o and falling temperature, there was a switch to subsolvus crystallisation of separate kand na-rich phases in the lujavrites. a continuum may have existed from temperatures >1100°c to increasingly low temperature (volatile-rich) magmas at c. 300°c before any discrete supercritical fluid phase separated. olivine compositions changed in the evolving melts, with the forsterite component approaching zero while the tephroite (mn) component increased, reaching a maximum of c. 16 mol % in the south qôroq complex (stephenson 1974). olivine eventually underwent reaction with melt, forming ironand sodium-rich amphibole. only in the kakortokite and lujavrite magmas did amphibole become a liquidus phase. magnetite was also lost by reaction with melt, producing aenigmatite (larsen & steenfelt 1974; larsen 1977). apatite underwent continuous changes becoming increasingly rich in sr and lree (p.g. hill, unpublished data) through a substitution dominated by ca2++p5+=ree3++si4+ with concomitant introduction of na, until vitusite (na3(ce,la,nd) (po4)2) with >20 wt% ree2o3 crystallised in the ilímaussaq agpaites (rønsbo et al. 1979; rønsbo 2008). finally, in the hyperagpaitic magmas phosphorus became mainly accommodated in the silicate-phosphate steenstrupine. after zr attained its maximum content in the magmas (c. 9000 ppm, bailey et al. 1981b), a separate zr mineral, eudialyte, joined the liquidus assemblage at the start of kakortokite crystallisation. thus, the magmas evolved over a crystallisation range of c. 800°c, from extremely simple, virtually monomineralic mineral-melt equilibria in the most primitive ygdc magma to extremely complex poly-component equilibria in the latest ilímaussaq residua. it may be envisaged that cumulate sequences were generated beneath the southern rift system at all depths at which magma batches underwent temporary or permanent residence. to produce the large volumes of salic differentiates at high levels, huge quantities of mafic/ultramafic cumulates must have been formed at depth as envisaged beneath the kenya rift (macdonald 2002). mantle sources there is general consensus that the gardar magmas, and specifically those contributing to the southern rift, originated in the mantle. crustal contamination was probably insignificant apart from the geographically and temporally restricted silica-oversaturated magma suites. because of the close affinity between older and younger gardar magmatic suites (e.g. motzfeldt and ilímaussaq or kûngnât and the ygdc) that succeeded each other over a period in excess of 100 ma, it has been argued that the magmas either originated in the lithosphere or that transient asthenospheric melts acquired lithospheric characteristics in the course of ascent (upton & emeleus 1987; macdonald & upton 1993; upton 1996). 110110 the relative poverty in normative diopside of the gardar basaltic compositions could be explicable through extensive high-pressure clinopyroxene fractionation. however, as the clinopyroxene deficiency that confers the troctolitic nature to the gabbros is present in both older and younger gardar mafic rocks it would be a remarkable coincidence if, over a time span of c. 140 ma, all the mafic magmas had undergone similar crystallisation histories. in the case of the nain province (labrador), morse (1982) considered it possible that the source of the anorthositic rocks was an unusually iron-rich mantle poor in clinopyroxene. for the gardar province it has been suggested that the mantle source was a metasomatised, clinopyroxene-poor lherzolite or even harzburgite (macdonald & upton 1993). the strongly fractionated ree patterns of the ygdc and other gardar basalts indicate that garnet was a residual phase during the melting processes. trace element and isotopic studies across the gardar province clearly demonstrate a heterogeneous mantle source. whilst basalts related to continental rifting (e.g. the east african rift, the basin and range province, oslofjorden and the carboniferous magmatism in south/central scotland) typically have ocean island basalt type (oib) incompatible element distributions, those of the gardar and specifically in the southern rift system differ in their higher k, sr, ba, p, f and lree contents and negative nb-ta anomalies, suggesting that supra-subduction zone metasomatism may have been involved, possibly dating back to palaeoproterozoic ketilidian events (goodenough et al. 2002). the high alkali and chlorine contents in the magmas may have resulted from interaction between carbonated peridotite and saline fluids or between peridotite and chloride-carbonate melts (klein-bendavid et al. 2009). these authors suggest that potassium may infiltrate peridotite during penetration of a saline component. the migration and focussing of such alkaline-halogen-rich fluids into the mantle wedge, supposed to have accompanied the southern gardar rift, could have been of prime importance in the processes leading, inter alia, to genesis of the ilímaussaq naujaites. it has been suggested for the kenya rift magmatism that there was interaction between a plume component (similar to an oib-source) and a heterogeneous lithospheric mantle with the qualification that the strong lithospheric signature makes identification of the plume component very difficult (macdonald 2003 and references therein). in the case of the eriksford formation lavas, it was noted by halama et al. (2002, 2004) that the trace element characteristics are comparable to those derived from oib-type sources. whatever the source, the paradox of the similarities between the older and younger gardar magmas remains to be resolved, despite the presumption of active asthenospheric convection during the long period separating them. the julianehåb batholith is considered to be of andean type, with the implication that it was a consequence of subduction at an ocean–continent plate boundary. the ketilidian orogeny took place at 1855–1723 ma (garde et al. 2002) and is considered to have involved oblique subduction of an oceanic plate subducting northwards beneath the craton margin (chadwick & garde 1996). the inferred enrichment of the sub-gardar mantle in k, ba, sr, p, f and lree, and negative nb-ta anomalies, was attributed to metasomatic modification by fluids or melts rising from the subducting oceanic plate (upton & emeleus 1987; macdonald & upton 1993; goodenough et al. 2002; marks et al. 2004; köhler et al. 2009). some of the gardar basaltic rocks are sufficiently potassic to justify use of the term shoshonitic (winther 1992; köhler et al. 2009). köhler et al. also report the high f content of gardar doleritic dykes and note that f-enrichment is also a characteristic of andean lavas, inferring that in both instances the element was derived from a subducting oceanic slab. the surviving ketilidian volcanic sequences on the craton north of the gardar province include tholeiitic pillow lava sequences several kilometres thick (higgins 1970; garde et al. 2002). they are characterised by low la/yb ratios suggesting that they represent high degrees of mantle melting (m. hamilton & b. upton, unpublished data, 2000). the lavas are as yet undated but are presumed to have an age of c. 2000 ma, and it is consequently proposed that their eruption left a restitic lithospheric mantle composed largely of refractory clinopyroxene-poor lherzolite or even harzburgite (upton 1996). fluids or hydrous silicic melts rising from the subducting oceanic plate may have first entered the overlying mantle wedge and, subsequently, the lithospheric mantle, or if the ketilidian lithosphere was sufficiently thick, entered directly into the lithospheric mantle. the only mantle xenoliths known from the gardar province are those in an aillikitic intrusion on illutalik, south-east of tuttutooq (described above). although these are severely deuterically altered, they contain glimmerite veins with high k, rb, ba, sr and lree contents (upton 1991). it is argued that metasomatism, initially through subduction-related processes, progressively changed refractory peridotitic rocks to more fusible compositions. accordingly a ‘chemical memory’, encapsulated at c. 1800 ma, 111 was not accessed until some 500–700 ma later when triggered during gardar cratogenesis (goodenough et al. 2002). the younger gardar activity involved great quantities of parental magmas. whilst all these were enriched, the extreme concentrations of incompatible elements at the ilímaussaq complex demand involvement of huge volumes of the mantle. it can only be surmised that processes of fractional melting, fluid transport and fractional crystallisation were capable of scavenging and concentrating trace components of the mantle on a very large scale. rifting of the columbia supercontinent the cratogenic gardar tectono-magmatism commenced between 1320 and 1280 ma. whilst the younger gardar rifting occurred between 1180 and 1140 ma, the 700 km long great abitibi dyke (canada), dated at 1140 ± 2 ma (krogh et al. 1987) is approximately colinear with the ygdc in reconstructions of pre-mesozoic greenland and labrador (macdonald & upton 1993). although sharing numerous features with the south greenland giant dykes, the great abitibi dyke has a tholeiitic composition (ernst & bell 1992). the ene–wsw-trending dyke extends south-west towards lake superior so that, if it was associated with the ygdc, the total extent would be c. 2000 km (fig. 121). from the great lakes southeastwards towards texas, the mid-continental rift is traceable a further c. 2000 km (hutchison et al. 1990). this rift, underlain by the largest bouguer anomaly on the craton, was associated with the keweenawan volcanism. the latter comprises a great volume of tholeiitic continental flood basalts that erupted from c. 1109 ma until c. 1086 ma (davis & paces 1990 and references therein). this volcanism has been attributed to adiabatic decompression of an asthenospheric mantle plume (nicholson & shirey 1990). collectively, the younger gardar, great abitibi and mid-continent rift events invite the speculation that each represented a component of rifting with concomitant basaltic magmatism across the columbia super-continent. rift propagation towards the southwest over a distance of some 4000 km may have resulted from intermittent lithospheric failure over some 70 ma. the gardar activity could represent an early stage in this process. topography of the younger gardar southern rift probably not more than 4 km of supercrustal cover have been stripped off the southern rift since it was an active volcanic rift zone. with regard to the surface topography we may envisage a stark, barren volcanic landscape within the columbia supercontinent that may have resembled the modern terranes of the danakil depression in ethiopia and the reykjanes peninsula in iceland, with parallel crater-chains, open fissures and normal fault escarpments marking the neovolcanic zone (fig. 122). such a landscape is envisaged to have formerly overlain the dyke swarm and associated basement shears of eastern tuttutooq (cf. fig. 12). the dykes of the main and igaliko swarms, which certainly reached shallow crustal levels, may have erupted relatively low-viscosity salic lavas across the rift, perhaps comparable to the kenya flood trachytes and phonolites. superimposed on this may have lain a chain of central volcanoes extending for some 60 km. the earliest of these, constructed above the narssaq and south qôroq complexes, would have been severely degraded by erosion and largely or wholly covered by younger extrusive rocks when the volcanoes over the tugtutôq, ilímaussaq and igdlerfigssalik complexes were active. possibly the tugtutôq central complex volcano and ilímaussaq were roughly coeval. the suggestion of an ilímaussaq volcano is contentious because there is consensus on the closed canada gad ? g na f greenland gd fig. 121. map of the great abitibi dyke (gad) and its possible relationship to the gardar giant dykes (gd). nipigon arm (na)of speculative triple junction. grenville front (gf). mid-continent rift and known extent of the keweenawan lavas indicated by grey ornamentation. modified from ernst & bell (1992). 112112 nature of the agpaitic magma chamber. the magma chamber rose high in the eriksfjord formation but there are no data for what might have lain above it, and it is conceivable that ilímaussaq crystallised within its own volcanic carapace. the relationship between sinistral faulting and dyke intrusion, together with the u-pb dating, suggests that ilímaussaq pre-dated the late stage igdlerfigssalik volcano. large linear volcanic systems such as boina and erta ale in ethiopia (barberi et al. 1970; barberi & varet 1970) and the harat khabar within the makkah-madinah-nafud volcanic lineament of western saudi arabia (camp et al. 1989) could serve as approximate models for the southern rift. erta ale comprises an elliptical structure c. 100 km long and 20–30 km broad in a region of rapid crustal extension along the median axis of the danakil depression (fig. 123). it displays evolution from simple fissural eruptions to complex central volcanoes with a generalised volumetric decrease in time from early transitional basalts through fe-rich intermediate compositions to highly differentiated products (trachytes and comendites). where silicic lavas were erupted as lava flows, the fissures are close to the central volcanoes. some of the trachytes of erta ale appear to have had high fluidity (barberi et al. 1970), much as is deduced for the gardar trachyte magmas. an evolution comparable to these ethiopian examples occurs in saudi arabia in the makkah-madina-nafud volcanic lineament. this extends for c. 600 km and has a sequence of vents that started with extensive extrusion of transitional olivine basalt that gave way to less voluminous flows including hawaiite, mugearite, benmoreite and trachyte. in the central vent area of harrat khaybar, the latest eruptions were of comendite (camp et al. 1989). it was suggested that primary mantle melts accumulated and evolved close to the crust–mantle boundary to the stage when they were copiously erupted along the whole volcanic lineament. in at least one case, some magma batches inferred to have been arrested in crustal fig. 122. lava fields of the western neovolcanic zone in south-west iceland. normal faults are prominent to the left. linear features to the right include crater chains and hyaloclastite ridges. the hengill central volcano is prominent in the far centre behind the steam columns from geothermal wells. photo by hjalti franzson. 113 reservoirs evolved further, resulting in comenditic residua (camp et al. 1989). many parallels may be drawn between the rift magmatism of these cenozoic instances and that envisaged for the younger gardar southern rift. if fissure eruptions accompanied the emplacement of the younger giant dyke and main dyke swarm, an extrusive carapace may have accreted above the tugtutôq complex, with progressively diminishing volumes of increasingly evolved lava over time in a manner comparable to these modern examples. if, as suggested, the tugtutôq complex underlay a rift-axial volcano, this is likely to have had a superstructure of quartz trachyte and alkali rhyolite extrusives. the ring dykes of the igdlerfigssalik complex and the central complex of tugtutôq suggest that any overlying volcanoes bore calderas. there is close petrological affinity between the intrusions of the southern rift and those of the kenya rift (macdonald & upton 1993). kenyan volcanoes that might serve as models include e.g. kilombe, suswa and longonot (fig. 124). 40° 44°42° 15° 13° 11° 9° l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l danakil horst red sea erta ale alayta boina et hi o pi an r if t djibouti somalian plateau et h io pi a n p la te au arabia l fig. 123. sketch map showing parts of ethiopia, the southern red sea and saudi arabia showing trends of rift faulting, spreading centres and the linear volcanic systems of boina and erta ale. modified from barberi et al. (1970). fig. 124. longonot volcano, kenya – a possible modern analogue for the late gardar volcanoes postulated for the south qôroq, igdlerfigssalik and tugtutôq complexes. the volcano rises c. 1000 m above the surrounding plains to a summit at 2776 m. 114114 summary younger gardar (1180–1140 ma) magmatic activity was principally manifested in two rift zones developed across the columbia supercontinent, viz. the northern (nunarsuit–isortoq) and the southern (tuttutooq–ilimmaasaq–narsarsuaq) rifts, in response to lithospheric extension. the tectono-magmatic evolution of the southern rift zone has been examined here. uplift and erosion of a few kilometres of cover have revealed a sequence of intrusions ranging from gabbros to highly evolved alkali granites and peralkaline nepheline syenites including agpaites. many of these intrusions are considered to have been related to surface volcanism. the model presented invokes transtensional movements occurring in conjunction with lithospheric attenuation and ascent of a narrow wedge of asthenospheric mantle. adiabatic melting of metasomatically modified lithospheric mantle within the rift zone is inferred as explanation for the unusual composition of the parental basaltic magmas. these had high al/ca ratios leading to crystallisation of troctolitic gabbros and anorthosites. the intrusions may conveniently be considered under three headings, namely giant dykes, dykes and stocks and ring dykes of central complexes. the intrusions in each of these categories are remarkable, if not for their size, shape or extent, then for many of the exceptional rock types that compose them, and unique in the case of ilímaussaq. giant dykes 200–800 m broad and smaller dykes <50 m broad dominated the early stages of the magmatic evolution, whilst central complexes characterised the later stages. the giant dykes and central complexes are largely composed of coarse-grained cumulates; the smaller dykes provide valuable petrographic and compositional data on magma types. lithospheric rupturing and emplacement of the older giant dyke complex (ogdc) marked the onset of activity. this intrusion comprises a near-complete spectrum of rock types from alkali gabbro via ferro-syenogabbros to syenites and peralkaline foyaites. after a time lapse of some tens of millions of years, a second and greater riftaxial rupturing event occurred, accompanied by intrusion of transitional basalt magma, to form the younger giant dyke complex (ygdc). this is considered to have arisen from a much more voluminous body of mafic magma most of which was retained in an underplated parental chamber near the moho. it is inferred to have resulted from a larger mantle melt fraction than that which had previously given rise to the ogdc and to have marked the acme of energy release and magma genesis related to the southern rift system. the ygdc fed an overlying lopolith at the unconformity between the palaeoproterozoic granites and the early gardar supracrustal strata. only relicts of this now remain. layered cumulate pods developed along the c. 145 km length of the ygdc may denote sites where vigorous convection was established. whilst most of the ygdc is composed of troctolite, differentiated products include peridotite, ferro-syenogabbro, syenite and both foyaite and quartz syenite to alkali granite. the closely related gabbroic to syenitic klokken complex to the south-east of the ygdc may be coeval and comagmatic with the ygdc. following the ygdc event, extensional energy was slowly dissipated, with intrusion of two remarkable dyke swarms. with time these show a tendency to diminish in width (and volume?) whilst increasing in degree of differentiation. dykes of the rift-axial main swarm are dominantly of hawaiite leading to trachyte and comendite. so-called big feldspar dykes (bfds) are important early components of the swarm and provide evidence of a deep crustal, compositionally stratified parental chamber in which hawaiitic magma became serially overlain by magma with compositions leading to quarz trachyte and comendite. the subsidiary igaliko dyke swarm, occurring to the south and east of the main dyke swarm, may have originated from smaller melt fractions of the mantle source. in this swarm phonolitic dykes are dominant amongst the salic members. anorthosite xenoliths in the troctolitic gabbros of the ygdc and klokken, in the doleritic component of the bfds, and elsewhere in the gardar province, indicate the presence of a large anorthositic body at depth. this body is considered to be cogenetic with the gardar alkaline magmas. the salic magmas of the province, including those of the southern rift magmatic system, are inferred to have been generated in a lower crustal chamber beneath a capping of anorthosite flotation cumulate. despite many features in common with the north american mid-proterozoic anorthosites, the gardar troctolites and anorthosites differ in being more potassic. in consequence, residual magmas followed alkaline lines of liquid descent. decline of fissuring and concomitant dyke intrusion coincided with rejuvenation of transcurrent (transform?) 115 fault systems oriented transverse to the rift. these, spaced c. 20–30 km apart, segment the rift zone with a total leftlateral offset of >18 km. weak spots provided by intersection of the faults and the rift axis offered potential access routes for magmas rising from the lower crust to shallow levels. these magmas are inferred to have arisen from the salic upper portions (grown, through maturation with time) of the stratified chambers mentioned above in relation to the bfds. the iron-rich, ferro-mugearitic magmas, inferentially generated at depth by massive plagioclase fractionation from relatively reduced melts, rarely, if ever, reached the shallow crust because of their high density. as the fe content of the magmas decreased through titanomagnetite fractionation, the benmoreite residues attained densities low enough for them to ascend by overhead stoping. benmoreitic magmas were the earliest to intrude at the ilímaussaq, tugtutôq and igdlerfigssalik complexes and were important in supplying dykes of the main swarm. the crustal weaknesses at the rift and fault intersections localised four of the central complexes of the southern rift. the two earliest of these, the narssaq and south qôroq complexes, were built up by successive magma batches ascending while the left-lateral faulting was still active. the narssaq complex consists of quartz syenite and alkali granite; the south qôroq complex consists of silica undersaturated products. since each intrusion at the south qôroq complex was of increasingly primitive magma, it provides confirmation for the presence of the compositionally stratified chamber inferred from the bfd evidence. as crustal equilibration took place following the slow demise of both fissuring and transcurrent faulting, late magma batches exploited the fault-controlled conduits to form the ilímaussaq and igdlerfigssalik complexes. location of the nearly contemporaneous tugtutôq central complex may have been dictated by the proximity of the older and younger giant dykes. whereas magma evolution primarily occurred along silica-undersaturated lines of descent there was exception to this in the central sector of the southern rift where benmoreite/trachyte residues evolved towards silicaoversaturated products. for a distance of c. 30 km, quartz syenites and alkali granites predominate in the tugtutôq central complex, the asorutit sector of the ygdc, the narssaq complex and the main dyke swarm. the occurrence of silica-oversaturated salic dykes beyond this sector is explicable by their lateral propagation. these more siliceous magmas probably originated from crustal contamination of their more mafic forerunners, brought about through further heating of the crust by hot fluids arising from the underlying mantle or by crystallisation of the underplated basaltic magma. despite the plethora of disparate and apparently unrelated rock types within the rift, a simple unitary genetic system is discernible, involving a bunch of closely related liquid lines of descent. the only magma types unrelated to this principal theme are those of the mela-aillikite–carbonatite association. it is suggested that these aberrant magmas resulted from rheomorphism of fusible masses of metasomites rich in diopside-phlogopite-apatite-calcite in the lithospheric mantle. the postulate of a mainly lithospheric mantle origin for all the gardar magmas is made to account for the close similarities between the older and younger gardar magmas. in view of the age difference of more than100 ma between them, this precludes a purely asthenospheric origin. although the world shows examples of a great many linear magmatic systems, the younger gardar southern rift system is unique with regard to its exposures, degree of preservation, layering features, and extreme and welldocumented compositional variations. acknowledgements i would like to dedicate this work to the memory of n.v. ussing and l.r. wager, knowing that ‘i ride upon the shoulders of giants’. i am grateful to geus for their support of this bulletin and specifically to l.m. larsen and a.a. garde for their indispensable help in its compilation. critical comments on the manuscript by j.c. bailey and t. andersen are gratefully acknowledged. my greatest debt, however, is to the former geological survey of greenland under the directorship of k. ellitsgaard-rasmussen for enabling most of the requisite field work to be undertaken. over the past 55 years, i have had the pleasure and benefit of collaborating with numerous scientists including a. berthelsen, a.b. blaxland, j. bondam, k. coe, j. craven, c.h. emeleus, j. ferguson, a.a. finch, j.g. fitton, i. gibson, k. goodenough, e.i. hamilton, w.h. harry, n. henriksen, j. köhler, l.m. larsen, r. macdonald, a. madsen, m. marks, a.r. martin, s.m. mingard, s. moorbath, r. nesbitt, i. parsons, n.j.g. pearce, j.d.a. piper, t.c.r. pulvertaft, h. scharbert, a. steenfelt, d. stephenson, j.w. stewart, h. sørensen, j.e. thomas, j.e. walton, b.j. watterson and w.s. watt. the carnegie trust for scottish universities provided financial assistance to travelling. 116116 references allaart, j.h. 1969: the chronology and petrography of the gardar dykes between igaliko fjord and redekammen, south greenland. rapport grønlands geologiske undersøgelse 25, 26 pp. andersen, s., bohse, h. & steenfelt, a. 1981a: a geological section through the southern part of the ilímaussaq intrusion. rapport grønlands geologiske undersøgelse 103, 39–42. andersen, s., bailey, j.c. & bohse, h. 1981b: zr-y-u stratigraphy of the kakortokite-lujavrite-sequence, southern ilímaussaq intrusion. rapport grønlands geologiske undersøgelse 103, 69–76. andersen, s., bohse, h. & steenfelt, a. 1988: the southern part of the ilímaussaq complex, south greenland 1: 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(ed.): layered intrusions, 331–363. elsevier science b.v. upton, b.g.j., emeleus, c.h., heaman, l.m., goodenough, k.m. & finch, a. 2003: magmatism of the mid-proterozoic gardar province, south greenland: chronology, petrogenesis and geological setting. lithos 68, 43–65. upton, b.g.j., craven, j.a. & kirstein, l.a. 2006: crystallisation of mela-aillikites of the narsaq region, gardar alkaline province, south greenland and relationships to other aillikitic-carbonatitic associations in the province. lithos 92, 300–319. upton, b.g.j., macdonald, r., odling, n., rämö t. & bagiński, b. 2013: kûngnât revisited. a review of five decades research into an alkaline complex in south greenland, with new trace element analyses and nd isotopic data. mineralogical magazine 77, 523–550. ussing, n.v. 1912: geology of the country around julianehaab, greenland. meddelelser om grønland 38, v–xi and 1–376. wager, l.r. & brown, g.m. 1968: layered igneous rocks, 588 pp. edinburgh: oliver & boyd ltd. wager, l.r. & deer, 1939: geological investigations in east greenland. part iii. the petrology of the skaergaard intrusion, kangerdlugssuak. meddelelser om grønland 105(4), 352 pp. wager, l.r., brown, g.m. & wadsworth, w.j. 1960: types of igneous cumulates. journal of petrology 1, 73–85. waight, t., baker, j. & willigers, b. 2002: rb isotope dilution analyses by mc-icpms using zr to correct for mass fractionation: towards improved rb-sr geochronology? chemical geology 186, 99–116. walker, g.p.l. 1993: basaltic-volcano systems. in: prichard, h.m. et al. (eds): magmatic processes and plate tectonics. geological society special publication (london) 76, 7–38. walton, b.j. 1965: sanerutian appinitic rocks and gardar dykes and diatremes north of narssarssuaq, south greenland. bulletin grønlands geologiske undersøgelse 57, 66 pp. (also meddelelser om grønland 179). watt, w.s. 1966: chemical analyses from the gardar igneous province, south greenland. rapport grønlands gelogiske under124124 søgelse 6, 92 pp. watt, w.s. 1968: petrology and geology of the precambrian gardar dykes on qaersuarsuk, south greenland. rapport grønlands geologiske undersøgelse 14, 51 pp. wegman, c.e. 1938: geological investigations in southern greenland. part 1. on the structural divisions of southern greenland. meddelelser om grønland 113, 148 pp. wiebe, r.a. 1992: proterozoic anorthosite complexes. in: condie, k.c. (ed.): proterozoic crustal evolution, 215–261. amsterdam: elsevier. winther, k.t. 1992: feldspar megacryst and anorthosite xenolithbearing dykes in the narssarssuaq area, south greenland. rapport grønlands geologiske undersøgelse 154, 49–59. woolley, a.r. 1987: lithosphere metasomatism and the petrogenesis of alkaline igneous rocks and carbonatites, malawi. journal of african earth science 6, 891–898. yoder, h.s. & tilley, c.e. 1962: origin of basalt magma: an experimental study of natural and synthetic rock systems. journal of petrology 3, 342–532. 125 126126 de nationale geologiske undersøgelser for danmark og grønland (geus) geological survey of denmark and greenland øster voldgade 10, dk-1350 copenhagen k denmark the series geological survey of denmark and greenland bulletin started in 2003 and replaced the two former bulletin series of the survey, viz. geology of greenland survey bulletin and geology of denmark survey bulletin. some of the twenty-one volumes published since 1997 in those two series are listed on the facing page. the present series, together with geological survey of denmark and greenland map series, now form the peer-reviewed scientific series of the survey. geological survey of denmark and greenland bulletin 1 the jurassic of denmark and greenland, 948 pp. (28 articles), 2003. edited by j.r. ineson & f. surlyk. 500.00 2 fish otoliths from the paleocene of denmark, 94 pp., 2003. by w. schwarzhans. 100.00 3 late quaternary environmental changes recorded in the danish marine molluscan faunas, 268 pp., 2004. by k.s. petersen. 200.00 4 review of survey activities 2003, 100 pp. (24 articles), 2004. edited by m. sønderholm & a.k. higgins. 180.00 5 the jurassic of north-east greenland, 112 pp. (7 articles), 2004. edited by l. stemmerik & s. stouge. 160.00 6 east greenland caledonides: stratigraphy, structure and geochronology, 93 pp. (6 articles), 2004. edited by a.k. higgins & f. kalsbeek. 160.00 7 review of survey activities 2004, 80 pp. (19 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pedersen, m. sønderholm, h.h. midtgaard, l.m. larsen, h. nøhr-hansen & a.k. pedersen. 300.00 20 review of survey activities 2009, 106 pp. (23 articles), 2010. edited by o. bennike, a.a. garde & w.s. watt. 220.00 21 exploration history and place names of northern east greenland, 368 pp., 2010. by a.k. higgins. 200.00 22 lithostratigraphy of the upper oligocene – miocene succession of denmark, 92 pp., 2010. by e.s. rasmussen, k. dybkjær & s. piasecki. 240.00 23 review of survey activities 2010, 84 pp. (19 articles), 2011. edited by o. bennike, a.a. garde & w.s. watt. 200.00 24 the east greenland rifted volcanic margin, 96 pp., 2011. by c.k. brooks. 200.00 25 upper cretaceous chalk facies and depositional history recorded in the mona-1 core, mona ridge, danish north sea. 2011. by k. anderskouv & f. surlyk. 200.00 127 26 review of survey activities 2011, 88 pp. (21 articles), 2012. edited by o. bennike, a.a. garde & w.s. watt. 200.00 27 neoglacial and historical glacier changes around 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higgins. 225.00 187 palynology and deposition in the wandel sea basin, eastern north greenland, 101 pp. (6 articles), 2000. edited by l. stemmerik. 160.00 188 the structure of the cretaceous–palaeogene sedimentary-volcanic area of svartenhuk halvø, central west greenland, 40 pp., 2000. by j. gutzon larsen & t.c.r. pulvertaft. 130.00 189 review of greenland activities 2000, 131 pp. (17 articles), 2001. edited by a.k. higgins & k. secher. 160.00 190 the ilímaussaq alkaline complex, south greenland: status of mineralogical research with new results, 167 pp. (19 articles), 2001. edited by h. sørensen. 160.00 191 review of greenland activities 2001, 161 pp. (20 articles), 2002. edited by a.k. higgins, k. secher & m. sønderholm. 200.00 geology of denmark survey bulletin (36–37; discontinued) 36 petroleum potential and depositional environments of middle jurassic coals and non-marine deposits, danish central graben, with special reference to the søgne basin, 78 pp., 1998. by h.i. petersen, j. andsbjerg, j.a. bojesen-koefoed, h.p. nytoft & p. rosenberg. 250.00 37 the selandian (paleocene) mollusc fauna from copenhagen, denmark: the poul harder 1920 collection, 85 pp., 2001. by k.i. schnetler. 150.00 prices are in danish kroner exclusive of local taxes, postage and handling note that information on the publications of the former geological survey of denmark and the former geological survey of greenland (amalgamated in 1995 to form the present geological survey of denmark and greenland) can be found on www.geus.dk http://www.geus.dk 128128 introduction uniqueness of the southern branch of the gardar rift nomenclature of place names general geological overview history of exploration gravity map the older giant dyke complex, tuttutooq marginal facies central facies the younger giant dyke complex composition of the magma crystallisation sequence internal structures the ygdc in the tuttutooq archipelago sissarluttooq marraat asorutit krydssø itillip saqqaa tripyramidal peak west-south-west of itillip saqqaa itillinnuujuk minor offshoots from the giant dykes narsaq gabbro and lopolithic relicts younger giant dyke extensions west and north of motzfeldt sø sydtungegletscher and syenitknold central complexes and late dykes klokken complex marginal gabbro syenogabbro and unlaminated syenite central layered series anorthosite xenoliths and plagioclase megacrysts in the ygdc and klokken gabbros origin of synformal layering in the younger giant dyke complex mela-aillikites, carbonate-silicate rocks and carbonatites mela-aillikite intrusions in the narsaq area and on tuttutooq mantle xenoliths diatremes other aillikite, carbonate-silicate and carbonatite dykes genesis of the ultramafic rocks narssaq complex south qôroq complex post-ygdc dyke swarms big feldspar dykes salic dykes of the main dyke swarm igaliko dyke swarm tugtutôq central complex mineralogy and geochemistry petrogenesis late basic dykes ilímaussaq complex augite syenite alkali granite and quartz syenite agpaitic syenites hidden layered series ilímaussaq parental magma micro-kakortokite dyke østfjordsdal syenite and igdlerfigssalik complex østfjordsdal syenite igdlerfigssalik complex the role of anorthosite laminated anorthosites genesis of the anorthosites emplacement mechanisms and tectonics emplacement mechanism of the giant dykes tectonics within the southern rift normal faulting transcurrent faulting evolution of the magmatic system of the younger gardar southern rift geochemical characteristics of the southern rift mafic magmas magma evolution in the southern rift zone magmatic differentiation in the lower crust magma chambers of the central complexes genesis of the silica-oversaturated magmas crystallisation histories mantle sources rifting of the columbia supercontinent topography of the younger gardar southern rift summary acknowledgements references geological survey of denmark and greenland bulletin 4, 2003, pp 49-52 49 for many years the existence of an oil-prone source rock off west greenland was challenged by industry. but since 1992 when active oil seeps were found onshore west greenland on the nuussuaq peninsula (fig. 1; christiansen et al. 1996; bojesen-koefoed et al. 1999), the question has changed focus to the age, distribution and potential of the source rock. five different oils – each with their own characteristics – have been reported by the geological survey of denmark and greenland (geus). one of these, a typical marine shalederived oil with a possible regional distribution, is known as the itilli oil. geochemical analysis suggests that it may have been generated from cenomanian–turonian age marine shales, equivalent to prolific source rocks known from ellesmere island, nunavut, canada. three of the other oils were generated from deltaic source rocks of albian, campanian and paleocene ages, while one is of unknown origin (bojesen-koefoed et al. 1999). the presence of a regional marine source rock is important to petroleum exploration; geus has therefore investigated the possible existence of mesozoic, in particular cenomanian–turonian, petroleum source rocks in west greenland offshore areas. since sediments older than the santonian are not known from any of the six wells drilled offshore west greenland (fig. 1), assessment of oil-prone source rocks in older sedimentary successions must rely on circumstantial evidence offered by oil chemistry data and analogy studies. petroleum in quantities amenable to chemical analysis has so far not been recovered from offshore. however, oilbearing fluid inclusions are known from the ikermiut-1 well (unpublished data 2001, phillips petroleum and geus), a gas-kick was recorded during drilling of the kangâmiut-1 well (bate 1997), and seismic data indicate hydrocarbons in many areas (cross-cutting reflectors, bright spots, smearing of seismic). petroleum exploration offshore west greenland suffered for many years under the misconception that oceanic crust covered vast areas, rendering the region unattractive. however, the presence of thick sedimentary successions and rotated fault blocks in cretaceous basins have been demonstrated to be present in areas previously believed to be underlain by cretaceous–tertiary oceanic crust (cf. chalmers & pulvertaft 2001). new high-quality seismic data, acquired by the seismic company tgs-nopec over recent years, combined with gravimetric data, have further demonstrated the presence of deep basins containing thick sedimentary successions in other areas (e.g. christiansen et al. 2002). despite the progress made over the past few years, the geological evolution of the davis strait region in general remains poorly understood, but new data on oil chemistry may shed some light on the history of this region. the itilli oil type based on its chemical characteristics, the itilli oil type is presumed to have been generated from marine shales of cenomanian–turonian age, but no known source rocks are fig. 1. map showing location of wells drilled offshore west greenland. the onshore oil seepage area in the disko–nuussuaq–svartenhuk halvø region is framed and includes the onshore exploration well gro#3. age of oils in west greenland: was there a mesozoic seaway between greenland and canada? jørgen a. bojesen-koefoed, hans peter nytoft and flemming g. christiansen geological survey of denmark and greenland bulletin 4, 49–52 (2004) © geus, 2004 exposed or have been drilled in west greenland. however, on ellesmere island, the lower part of the kanguk formation comprises excellent, highly oil-prone marine shale source rocks of presumed cenomanian–turonian age (núñezbetelu 1993). the deposits are, however, thermally immature and cannot be directly compared to the itilli oil from onshore west greenland more than 1000 km to the south. however, artificial maturation by hydrous pyrolysis of samples of the kanguk formation generates bitumen that shares a number of important characteristics with the itilli oil type. these characteristics include: light-end skewed n-alkane distribution, pristane/phytane ratio less than 2, abundant tricyclic triterpanes, abundant 28,30-bisnorhopane and diasteranes, plus a predominance of c27 steranes while maintaining relatively high proportions of c28 steranes (fig. 2). a mesozoic seaway between greenland and canada? the recent demonstration of more or less continuous deep sedimentary basins offshore west greenland (christiansen et al. 2002), the occurrence of reworked marine upper jurassic palynomorphs in the qulleq-1 well (christiansen et al. 2001) together with the apparent relationship between the itilli oil and pyrolysates of the lower kanguk formation, open possibilities for the existence of regional mesozoic marine source rocks in the offshore areas. this implies the existence of a seaway between greenland and canada during the mesozoic, connecting the proto-atlantic to the proto-arctic ocean – a palaeo-davis strait, partly analogous to the ‘cretaceous western interior seaway’ (cwis) of north america (caldwell & kauffman 1993; dean & arthur 1998). the cwis developed as a foreland basin east of the rising cordillera along the western margin of the north american continent, whereas the davis strait is a product of rifting and strike-slip movements, albeit without sea-floor spreading. this is noteworthy since many published palaeogeographic maps feature a spreading ridge along the axial part of the davis strait. despite the differences in origin, the morphological analogy between a palaeo-davis strait and the cwis is clear and the potential for deposition of similar types of sediments in both settings exists. the evolution of the cwis is relatively well constrained, and petroleum accumulations, derived from cenomanian–turonian age marine source rocks, are known more or less throughout the entire extent of the seaway – from the southern part of the usa to the canadian arctic. a series of marine oils of cenomanian–turonian age from the cwis, plus a number of other oils (table 1), have been collected and analysed to serve as a reference for comparison with itilli oils from west 50 fig. 2. characteristics of the itilli oil type (modified from bojesen-koefoed et al. 1999). gas chromatogram shows light-end skewed n-alkane distribution and pristane/phytane < 2 (numbers: n-alkane carbon number, a: pristane, b: phytane). triterpanes monitored by the m/z 191 fragmentogram show abundant tricyclics (t23: c23 tricyclic triterpane) and notable proportions of 28,30-bisnorhopane (h28). ts: trisnorneohopane; tm: trisnorhopane; h29: norhopane; h30: hopane; h33: trishomohopane. steranes monitored by the m/z 217 and m/z 218 fragmentograms show a high abundance of diasteranes (d27) at moderate levels of thermal maturity indicated by c29 sterane s/(s+r) epimerisation ratio of approximately 0.50 (c29 sterane: s29; s: 20s epimer; r: 20r epimer), and a relatively high abundance of c28 steranes (s28) compared to c27 (s27), c29 (s29) and c30 (s30) steranes. greenland. in addition to cenomanian–turonian derived marine oils from the cwis, the reference sample database includes oils generated from upper jurassic source rocks in the jeanne d’arc basin (newfoundland, eastern canada) and in the north sea, the cretaceous age heron h-73 oil from offshore eastern canada, and the cambro-ordovician shoal point crude from western newfoundland (fig. 3). the distribution of a series of diatom-derived oil constituents known as 24-norcholestanes has proven to be agediagnostic, and a standard plot for assessment of oil source rock maximum age has been devised by holba et al. (1998). using this plot, the reference samples show an n–s trend among the cwis oils of cenomanian–turonian age (fig. 4). the single cretaceous oil from offshore eastern canada plots between the denver and alberta basin oils, as expected from its geographical position. upper jurassic oils from the jeanne d’arc basin and the north sea groups occur in a narrow band at the expected position, whereas the cambro-ordovician oil yields a ‘palaeozoic’ age. superimposing data from west greenland itilli oils onto the reference oil plot shows that samples in which admixture of oil from other sources can be recognised all yield rather young ages, whereas pure itilli oil samples show cretaceous or even late jurassic source rock ages (fig. 4). hence, age-diagnostic biological marker data support the existence of a regional cretaceous age marine petroleum source rock, in addition perhaps to an upper jurassic source rock. a characteristic feature of cenomanian–turonian age oils from the canadian arctic is a relatively high abundance of c28 regular steranes, compared to c27 and c29 regular steranes, whereas upper jurassic oils generally show a rather low abundance of c28 regular steranes. some itilli oils from west greenland show sterane distributions very similar to cenomanian–turonian age oils from the canadian arctic, others may show distributions rather similar to upper jurassic oils, and some may show intermediate distributions. hence, regular sterane data further support the notion of a cretaceous plus perhaps an additional upper jurassic marine petroleum source rock in the davis strait region, and thus the existence of a mesozoic seaway between greenland and canada. conclusions the itilli oil type from onshore central west greenland is an oil derived from marine shale, and shows clear similarities to pyrolysates of immature cenomanian–turonian age oilprone source rocks from the canadian arctic. analysis of age-diagnostic biological markers and the distribution of regular steranes indicate a source rock age similar to that of cenomanian–turonian age oils from the canadian arctic or of older upper jurassic oils from the jeanne d’arc basin. geochemical data, combined with other indications of petroleum in west greenland offshore areas, support the existence of one or more mesozoic marine petroleum source rocks in the larger davis strait area, and hence the notion of a mesozoic seaway between greenland and canada. acknowledgements analyses of samples were financed by the bureau of minerals and petroleum, government of greenland (centur project, grant no. 69.41.06). samples provided by the united states geological survey (denver), the geological survey of canada (calgary and halifax), and the canadian new-foundland offshore petroleum board (st. john’s) are gratefully acknowledged. 51 fig. 3. cretaceous western interior seaway (cwis) of north america and the north atlantic. red dots: approximate positions of analysed oil samples. references bate, k.j. 1997: interpretation of the basal section of well kangâmiut-1, offshore southern west greenland. danmarks og grønlands geologiske undersøgelse rapport 1997/76, 24 pp. bojesen-koefoed, j.a., christiansen, f.g., nytoft, h.p. & pedersen, a.k. 1999: oil seepage onshore west greenland: evidence of multiple source rocks and oil mixing. in: fleet, a.j. & boldy, s.a.r. (eds): petroleum geology of northwest europe: proceedings of the 5th conference, 305–314. london: geological society. caldwell, w.g.e. & kauffman, e.g. 1993: evolution of the western interior basin. geological association of canada special paper 39, 679 pp. chalmers, j.a. & pulvertaft, t.c.r. 2001: development of the continental margins of the labrador sea – a review. in: wilson, r.c.l. et al. (eds): non-volcanic rifting of continental margins: a comparison of evidence from land and sea. geological society special publication (london) 187, 77–105. christiansen, f.g., bojesen-koefoed, j.a., dam, g., nytoft, h.p., larsen, l.m., pedersen, a.k. & pulvertaft, t.c.r. 1996: the marraat oil discovery on nuussuaq, west greenland: evidence for a latest cretaceous – earliest tertiary oil prone source rock in the labrador sea – melville bay region. bulletin of canadian petroleum geology 44, 39–54. christiansen, f.g. et al. 2001: petroleum geological activities in west greenland in 2000. geology of greenland survey bulletin 189, 24–33. christiansen, f.g., bojesen-koefoed, j.a., chalmers, j.a., dalhoff, f., marcussen, c., nielsen, t., nøhr-hansen, h. & sønderholm, m. 2002: petroleum geological activities in west greenland in 2001. geology of greenland survey bulletin 191, 84–89. dean, w.e. & arthur, m.a. 1998: stratigraphy and paleoenvironments of the cretaceous western interior seaway, usa. sepm concepts in sedimentology and paleontology 6, 255 pp. holba, a.g., dzou, l.i.p., masterson, w.d., hughes, w.b., huizinga, b.j., singletary, m.s., moldowan, j.m., mello, m.r. & tegelaar, e. 1998: application of 24-norcholestanes for constraining source age of petroleum. organic geochemistry 29, 1269–1283. núñez-betelu, l.k. 1993: rock-eval/toc pyrolysis data from the kanguk formation (upper cretaceous), axel heiberg and ellesmere islands, canadian arctic. geological survey of canada open file 2727, 30 pp. 52 fig. 4. standard diagram for assessment of oil source rock maximum age using the nordiacholestane ratio (modified from holba et al. 1998). note n–s trend among cenomanian–turonian from the cretaceous western interior sea-way. mixed oils yield young source rock ages whereas pure itilli oils show cretaceous or even jurassic source rock ages. wcsb: western canadian sedimentary basin. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 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<feff4f7f752890194e9b8a2d5b9a5efa7acb76840020005000440046002065874ef65305542b8f039ad876845f7150cf89e367905ea6ff0c4fbf65bc63d066075217537054c18cea3002005000440046002065874ef653ef4ee54f7f75280020004100630072006f0062006100740020548c002000520065006100640065007200200035002e0030002053ca66f465b07248672c4f86958b555f3002> >> >> setdistillerparams << /hwresolution [2400 2400] /pagesize [595.000 842.000] >> setpagedevice geological survey of denmark and greenland bulletin 7, 2004, p 57-60 57 the geological survey of denmark and greenland (geus) and the bureau of minerals and petroleum (bmp, under the government of greenland) have co-operated on the international promotion of the mineral resources of greenland for more than ten years. the government of greenland follows a strategy aimed at the development of a mining and petroleum sector in greenland capable of yielding a significant proportion of the national income. to reach this goal it is necessary to attract international investment. in respect of mineral exploration, many parts of greenland can still be considered virgin territory and it is therefore vital that all data relevant for the identification of possible exploration targets are available to the international mining industry. geus has produced many compilations of geoscience data for that purpose in traditional reports, on cd-roms and in scientific journals. in 2004, a new source of geoscience information was developed based on an interactive gis facility on the internet, and mineral exploration data and information from a region in central west greenland are now accessible at the greenland mineral occurrence map (gmom) website at geus (fig. 1; www.geus.dk/gmom). technically, this new facility will be maintained and developed in accordance with general principles for internet services adopted by geus (e.g. tulstrup 2004). new information from other regions of greenland will gradually be added. compilation of available information in databases and company reports for the specific purpose of producing a completely up-todate mineral occurrence map in electronic form on the internet, existing data and information in the databases and archives of geus have been assessed, and the relevant extracts and syntheses are included in a new database develon-line presentation of mineral occurrences in greenland leif thorning, lisbeth aa. christensen, bo møller nielsen, frands schjøth and henrik stendal geological survey of denmark and greenland bulletin 7, 57–60 (2005) © geus, 2005 fig. 1. the greenland mineral occurrence map (gmom) is a new on-line service to the international mining industry introduced in january 2005. the figure shows the gmom website with an index map of greenland and various explanatory notes. the first region to be included in the map is part of west greenland. oped as part of the gimmex system (thorning et al. 2004). for each locality the database contains a compiled summary of relevant data, including geus’ assessment of the site and its relevance for a general understanding of the economic potential of a region. the database is used for the compilation and assessment process in geus in preparation for the production of the map. a selection of attributes, i.e. properties related to each locality, is used for the construction of the maps in the esri arcgis environment. the selected attributes are available as tables in the gis environment, giving the user of the gmom website access to data and information required for further study of the mineralised locality. among the attributes are geus’ suggestions for a likely standard model for the mineralisation according to eckstrand et al. (1996). the importance of the locality is indicated by a first, rough classification as to whether the locality is an indication, a showing, a prospect or a deposit, and a code signifying its classification according to unece (1997). the database also contains standardised descriptions of the localities, including photographs, detailed maps and references when available. some localities of similar nature have been grouped, and one locality has been chosen as type locality for the group. the content of the gmom database will increase as additional localities from other regions of greenland are included. new data will originate from geus’ own investigations and from company exploration activities as these become public. once in the database they will be available for inclusion in the gmom, which thus becomes a dynamic map display of exploration data and information. gis interface to on-line greenland mineral occurrence map the gmom database is the source of information for the on-line greenland mineral occurrence map. this on-line facility has been created using the esri arcims environment and users only need their internet browser to gain access to the greenland mineral occurrence map. apart from the map itself, the gmom website contains an introduction to the map, and provides a brief background together with explanations and definitions, references and other relevant information. from the index map at the gmom website, the user can choose the area of interest among pre-set options. the map server controls the session in a new window and places a fully operational gis system at the user’s disposal (fig. 2). the information from the gmom database is the primary data to be displayed. however, a number of optional gis layers are available for inclusion on the map according to the user’s choice. based on data from other gimmex geoscience databases at geus, different backdrop maps for the display of mineralised sites are available, such as a topographic map, on which features such as rivers and lakes can be shown or omitted (fig. 2), a digital elevation model (fig. 3) and coloured anomaly maps of airborne geophysical data or geochemical data (fig. 4). the mineralised localities are shown as coloured symbols, which correspond to different commodity groups, such as precious metals, base metals, industrial minerals, etc., while the number is the locality identification in the gmom database. 58 fig. 2. opening view of gmom, showing the main part of the first region in greenland to be presented. various backdrop maps can be used. data from the mineralised localities can be inspected using the i-button or the hyperlink-button (the flash). the arcims standard gis tools available allow zooming, selection of layers, searches in the table of the active layer, measuring distances on the map and displaying or omitting features. selections of sites can be made according to name, location, commodity or any of the other attributes stored in the gmom database. the properties of individual localities can be obtained through selection of a location symbol on the map. clicking on the symbol with the mineral occurrence layer selected as active will bring up a table of information in the frame below the map with all the attributes as defined on the gmom website. this includes a hyperlink to a file in pdf format with a description of the mineralisation and its surroundings. the description can be called to the screen by using the hyperlink in the table, or by using the gis hyperlink button (the flash) before clicking on the symbol for the locality on the map. the user can select different backdrops for the map from the choices available and customise the map to his purpose. since this is all done in an arcims environ59 fig. 3. example from gmom showing the use of a digital elevation model as backdrop for the localities, which are coloured according to the type of commodity. the table displayed in the lower frame shows information for locality number 150. fig. 4. inspection of the summary information sheet of locality number 94 in the gmom database, a probable copper showing. the backdrop map shows a total intensity magnetic field anomaly map (shaded relief). the description of the site can also be downloaded as a pdf file (separate window). 60 ment, map and attribute data can be downloaded to the user’s own environment in a number of ways and incorporated with the user’s own data as required. at the launch date (january 2005) the on-line segment of the gmom covers systematically arranged data and information for a region of west greenland between latitudes 66° and 70°15´n, relying mostly on the recently completed mineral resource assessment programme for central west greenland (stendal et al. 2004). this includes much new data from 160 numbered localities within the area. however, there may be more interesting localities than those presently compiled since some of the areas within the region are still only explored at reconnaissance level. this illustrates the advantage of presenting the greenland mineral occurrence map in a dynamic form on the internet, rather than printed on paper. new information will be entered into the database as it becomes available from released company data or from further geus activities, and up-todate versions of the maps will therefore be available to users in the future. more data to come in 2005 more regional data from mainly south greenland and north-east greenland will be included in the greenland mineral occurrence map. furthermore, thematic maps for the entire country will be built into the system, and other accessory information will be made available for professional users interested in the mineral exploration potential of greenland. references eckstrand, o.r., sinclair, w.d. & thorpe, r.i. (eds) 1996: geology of canadian mineral deposit types. geology of canada 8, 640 pp. ottawa: geological survey of canada. stendal, h., nielsen, b.m., secher, k. & steenfelt, a. 2004: mineral resources of the precambrian shield of central west greenland (66° to 70°15´n). part 2. mineral occurrences. danmarks og grønlands geologiske undersøgelse rapport 2004/20, 212 pp. thorning, l., christensen, l. schjøth, f. & stendal, h. 2004: greenland mineral occurrence map. danmarks og grønlands geologiske undersøgelse rapport 2004/28, 52 pp. tulstrup, j. 2004: environmental data and the internet: openness and digital data management. geological survey of denmark and greenland bulletin 4, 45–48. unece, 1997: united nations international framework classification for reserves/resources. solid fuels and mineral commodities. united nations economic commission for europe energy/wp.1/r.77, 174 pp. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: lth@geus.dk gelogical survey of denmark and greenland bulletin 4,2003, pp 37-40 37 the danish nationwide aquatic monitoring programme (kristensen et al. 1991; kronvang et al. 1996) is one of the most comprehensive in europe and, compared with other national programmes, was put in place at an early date: the systematic monitoring of freshwater lakes began in 1989. as with all such programmes of long-term environmental monitoring, its value multiplies with time as an increasing number of scientific hypotheses can be tested and management questions addressed using the growing body of data. however, the aquatic monitoring data, as for many natural systems, are limited to the scale of decades and so may not identify important longer-term trends or be useful in the determination of ‘baseline’ (pre-industrial or pre-agricultural impact) conditions. the study of lake sediments, palaeolimnology, offers a means of extending the monitoring of lake status back in time and can be used to determine baseline conditions (e.g. for water chemistry variables), to define trends and rates of change, and to identify the causes of change (battarbee 1999). the data obtained from palaeolimnological studies can provide a yardstick against which current and future changes may be assessed. furthermore, restoration targets for lakes impacted by for example enhanced nutrient loading, will be aided by a knowledge of the environmental conditions of a pre-disturbance state, defining both potential targets for restoration and the limits to what a restoration project might achieve (battarbee 1999). application of palaeolimnology to the eu’s water framework directive palaeoecology has been a core area of research at the geological survey of denmark and greenland (geus), going back to the earliest days of the geological survey of denmark (rasmussen 1988). this long tradition has been followed and developed by successive research teams with new skills and techniques being taken on board. for example, advances in numerical techniques in the 1980s led to the development of transfer function methods for estimating past water chemistry variables from the microfossil record (birks et al. 1990). over the last 13 years, a transfer function to estimate past in-lake nutrient (total phosphorus, tp) concentrations for danish lakes from sedimentary diatom assemblages has been created, developed and applied in many studies (e.g. anderson & odgaard 1994; bennion et al. 1996; bradshaw 2001; bradshaw et al. 2002). the wealth of palaeolimnological data on numerous freshwater and brackish lakes in denmark, accumulated over many years, provides a crucial resource relevant to the implementation of the european union’s water framework directive. this directive, from 2000, aims to achieve a ‘good’ ecological status for all eu water bodies by the year 2015. one of the key considerations of the water framework directive is the determination of baseline conditions, or ‘reference status’, the yardstick against which contemporary water quality status can be assessed and future targets can be realistically defined. the reference status need not, of course, be a stable condition, nor will the unimpacted status be the same for different lakes in different landscapes (søndergaard et al. 2003). using the geological record to assess the changing status of danish lakes emily g. bradshaw and peter rasmussen fig. 1. map of denmark with the 21 lakes (blue dots) included in additional palaeolimnological analyses made for the european union’s water framework directive report (amsinck et al. 2003) and other lakes (red dots) mentioned in the text. geological survey of denmark and greenland bulletin 4, 37–40 (2004) © geus, 2004 two recent reports (amsinck et al. 2003; søndergaard et al. 2003) that are the result of collaboration between the national environmental research institute (neri), the danish forest and nature agency, the danish counties and the survey, set the foundations for implementation of the eu’s water framework directive for danish lakes. the archive of palaeolimnological data collated by the survey over many years, and complemented by parallel studies at neri, proved an invaluable basis for the consideration of reference status for danish lakes. in addition, new analyses (diatoms, geus; cladocerans, neri) were made for 21 lakes (fig. 1) for four time-slices: 1850, 1900, 1950 and 2000. this study was made possible, on a very limited budget, due to the wellkept archive of lake sediments held by geus: lake sediment cores that had been sampled and dated for a previous pollenbased study (nielsen 2003) were made available for additional analyses. lake eutrophication – a long-term view in addition to the survey’s long tradition of pollen analysis, many lake sediment-based studies have utilised the techniques of diatom and plant macrofossil analyses. the example given here (fig. 2) shows data from søgård sø. on the basis of the macrofossil data for aquatic plants, the record has been divided into three stages. the lower section of the sediment core (c. 200 to 147 cm) records sparse remains of submerged macrophytes, but a relatively well-developed community of floating-leaved plants (e.g. nymphaea: waterlily). in this earlier part of the record, the diatom assemblage reflects high total phosphorus concentrations in the lake. the second stage (c. 147 to 78 cm) shows a much expanded vegetation of submerged aquatic macrophytes (e.g. stratiotes aloides: water-soldier; nitella and chara: stoneworts; potamogeton: pondweed). egg-cocoons of the leech piscicola geometra also indicate the presence of submerged macrophytes at this time (odgaard & rasmussen 2001). the third stage (c. 78 to 0 cm) sees a greatly reduced number of plant macrofossil remains and p. geometra cocoons in the sediment, coinciding with a major increase in the diatom-inferred tp concentration. the record from søgård sø demonstrates the pattern of eutrophication seen in many danish lakes during the last century. before 1940 the lake had a rich aquatic macrophyte vegetation, but with increased nutrient concentrations in the 1940s and later, this macrophyte cover disappeared. this example is from just one of many danish lakes that have been the subject of palaeolimnological studies. figure 3 summarises some of the results from the study of 21 danish lakes carried out in connection with the water framework directive report (amsinck et al. 2003). the results of the analysis of sedimentary diatom remains suggest that many of the lakes underwent eutrophication during the past 150 years, but that in-lake nutrient concentrations were already surprisingly high for many sites in 1850. this conclusion is supported by the few studies of longer lake sediment sequences for danish lakes, e.g. langesø (anderson & odgaard 1994), dallund sø (bradshaw 2001), gudme sø and sarup sø (unpublished data, p. rasmussen and e.g. bradshaw) that demonstrate that some lakes have been impacted by human activities over the scale of centuries and millennia. thus, defining a true baseline condition for many lakes 38 fig. 2. summary stratigraphic data for søgaard sø (unpublished data, p. rasmussen and n.j. anderson). requires the consideration of long time periods, only obtainable by use of the geological record. environmental management requires both foresight and geological perspectives the study of palaeolimnological records has much to contribute to lake management questions and problems. as well as determining when and to what extent danish lakes were influenced by human activities, sites may be identified that have experienced relatively little impact over time and thus may be of greater conservation value (again, providing the yardstick against which to judge contemporary systems). also, rates of change can be informative with respect to the resistance of ecosystems to forcing events and the likelihood of future stability. pressing environmental problems and management questions, for example those posed by the eu water framework directive, must draw on knowledge and experience gained over many years. the danish nationwide aquatic monitoring programme has provided a rich resource to draw on when addressing the requirements of this directive, putting denmark in a stronger position than many other european countries. the long tradition of environmental history research in denmark, at the survey in particular, has likewise provided a strong base of knowledge, the fruits of which continue to be harvested, as illustrated by the response to the demands of the water framework directive. the management questions and the demands of the danish society in the future may not be predictable today (e.g. problems relating to climate change). therefore, the importance of maintaining relevant and well-established monitoring programmes for natural systems cannot be overstated (see also sand-jensen 1997; wolf 2003). their continuation, however, requires foresight and a belief in investment for the future. another illustration of this comes from the area of forest history research. since 1948, survey scientists have monitored natural forest development in denmark, one of the longest surveys of its kind in europe. the experience and knowledge gained has been of vital importance for the understanding of natural forest processes and for nature conservation in danish forests. once interrupted or discontinued, the value of a monitoring programme diminishes greatly. in the same way, research expertise developed over many years is a valuable resource that grows over time to become more than the sum of a research group’s parts. without the foresight to support and continue relevant and successful programmes, the knowledge-base that has been built up can rapidly disintegrate, but the effects of this may only be fully appreciated in subsequent decades (see e.g. sand-jensen 1995). of human history, anne knudsen, chief editor of the national danish newspaper weekendavisen, has incisively stated that: “lacking knowledge of history leaves us vulnerable and wide open to cock-and-bull stories and demagogy”. the same can also certainly be said for environmental history. for example, based on limited information, it is a commonly held belief that the pollution of danish lakes is a relatively modern phenomenon, associated with post-industrial population growth and agricultural development. the study of environmental history shows us that this is a misconception. references amsinck, s.l. et al. 2003: vandrammedirektivet og danske søer. del 2: palæoøkologiske undersøgelser. faglig rapport fra danmarks miljøundersøgelser 476, 120 pp. anderson, n.j. & odgaard, b.v. 1994: recent palaeolimnology of three shallow danish lakes. hydrobiologia 256–257, 411–422. battarbee, r.w. 1999: the importance of palaeolimnology to lake restoration. hydrobiologia 395–396, 149–159. bennion h., juggins s. & anderson, n.j. 1996: predicting epilimnetic phosphorus concentrations using an improved diatom-based transfer function and its application to lake eutrophication management. environmental science and technology 30, 2004–2007. birks, h.j.b., line, m., juggins, s., stevenson, a.c. & ter braak, c.j.f. 1990: diatoms and ph reconstruction. philosophical transactions of the royal society of london, series b – biological sciences 327, 263–278. bradshaw, e.g. 2001: linking land and lake. the response of lake nutrient regimes and diatoms to long-term land-use change in denmark, 118 pp. unpublished ph.d. thesis, university of copenhagen, denmark. bradshaw, e.g., anderson, n.j., jensen, j.p. & jeppesen, e. 2002: phosphorus dynamics in danish lakes and the implications for diatom ecology and palaeoecology. freshwater biology 47, 1963–1975. 39 fig. 3. mean diatom-inferred total phosphorus values (di-tp) for the additional sites studied for the water framework directive. maximum and minimum values are shown as ‘error’ bars. (note: n=17; the estimation of tp concentrations was not possible for all sites, some being outside the range of the transfer function model.) 40 kristensen, p., jensen, j.p., jeppesen, e. & erlandsen, m. 1991: ferske vandområder – søer. vandmiljøplanens overvågningsprogram 1990. faglig rapport fra danmarks miljøundersøgelser 38, 104 pp. kronvang, b, svendsen, l.m., larsen, s.e. & jensen, j.p. 1996: monitoring and modelling of nutrient loads in danish streams and lakes. in: riverine input to coastal areas – notes from a workshop on methodology, tema nord 1996: 529, 53–60. copenhagen: nordic council of ministers. nielsen, a.b. 2003: pollen based quantitative estimation of land cover. relationships between pollen sedimentation in lakes and land cover as seen on historical maps in denmark ad 1800, 135 pp. unpublished ph.d. thesis, university of copenhagen, denmark. odgaard, b.v. & rasmussen, p. 2001: the occurrence of egg-cocoons of the leech piscicola geometra (l.) in recent lake sediments and their relationship with remains of submerged macrophytes. archiv für hydrobiologie 152 (4), 671–686. rasmussen, l.b. 1988: en jordisk krønike. træk af dgu’s historie 1888–1988, 114 pp. københavn: danmarks geologiske undersøgelse, miljøministeriet. sand-jensen, k. 1995: furesøen gennem 100 år. naturens verden 5, 176–187. sand-jensen, k. 1997: eutrophication and plant communities in lake fure during 100 years. in: sand-jensen, k. & pedersen, o. (eds): freshwater biology. priorities and development in danish research, 26–38. copenhagen: gad. søndergaard, m., jeppesen, e., jensen, j.p., bradshaw, e., skovgaard, h. & grünfeld, s. 2003: vandrammedirektivet og danske søer. del 1: søtyper, referencetilstand og økologiske kvalitetsklasser. faglig rapport fra danmarks miljøundersøgelser 475, 142 pp. wolf, a. 2003: tree dynamics in draved forest. a long-term study of a temperate deciduous forest in denmark, 164 pp. unpublished ph.d. thesis, the royal veterinary and agricultural university, copenhagen, denmark. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: egb@geus.dk << /ascii85encodepages false /allowtransparency false /autopositionepsfiles true /autorotatepages /all /binding /left /calgrayprofile (dot gain 20%) /calrgbprofile (srgb iec61966-2.1) /calcmykprofile (u.s. web coated \050swop\051 v2) /srgbprofile (srgb iec61966-2.1) /cannotembedfontpolicy /warning 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<feff4f7f752890194e9b8a2d5b9a5efa7acb76840020005000440046002065874ef65305542b8f039ad876845f7150cf89e367905ea6ff0c4fbf65bc63d066075217537054c18cea3002005000440046002065874ef653ef4ee54f7f75280020004100630072006f0062006100740020548c002000520065006100640065007200200035002e0030002053ca66f465b07248672c4f86958b555f3002> >> >> setdistillerparams << /hwresolution [2400 2400] /pagesize [595.000 842.000] >> setpagedevice geological survey of denmark and greenland bulletin 20, 2010, 1–8 1 geological survey of denmark and greenland bulletin 20 • 2010 review of survey activities 2009 edited by ole bennike, adam a. garde and w. stuart watt geological survey of denmark and greenland ministry of climate and energy 22 geological survey of denmark and greenland bulletin 20 keywords geological survey of denmark and greenland, survey organisations, current research, denmark, greenland. cover photographs from left to right 1. the swedish ice breaker oden at the north pole during the lomrog ii cruise on 22 august 2009. photo: adam jeppesen. 2. folded permian volcanic rocks in vietnam. photo: stig a. schack pedersen. 3. field work in west greenland. photo: denis schlatter. 4. a seismometer is placed in a pit on the greenland ice sheet. photo: hans thybo. frontispiece: facing page tanzanian woman using mercury to extract gold. photo: peter w.u. appel. chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, department of earth sciences, university of aarhus; minik rosing, geological museum, university of copenhagen; finn surlyk, department of geography and geology, university of copenhagen scientific editors: ole bennike, adam a. garde and w. stuart watt editorial secretaries: jane holst and esben w. glendal referees: (dk = denmark etc.; numbers refer to first page of reviewed article): anonymous (43, 83, 99); james d. appleton, uk (87); stefan bachu, ca (95); terje bjerkgård, no (59); didier bonijoly, fr (95); gregers dam, dk (75); david lundbek egholm, dk (55); ida fabricius, dk (15, 47); henrik friis, dk (47); ole graversen, dk (99); robert hall, gb (91); claus heinberg, dk (75); karin högdahl, se (71); cecilia jelinek, se (39); birthe eg jordt, dk (27); ralf klingel, de (39); karen luise knudsen, dk (35); john a. korstgård, dk (51); gunnar larsen, dk (31); nicolaj krog larsen, se (31); kaj lax, se (103); jan mangerud, no (35); christoph mayer, de (55); sebastian mernild, us (79); john myers, au (67); allan aasbjerg nielsen, dk (71); ole bjørslev nielsen, dk (23); poul østergaard, dk (19); heikki papunen, fr (63); christophe pascal, no (19); asger ken pedersen, dk (59); gunver krarup pedersen, dk (23); iain pitcairn, se (63); martin sønderholm, dk (51); jette sørensen, dk (27); henrik stendal, gl (87); holger stünitz, no (67); ole v. vejbæk, dk (15); ian watkinson, gb (91); jacob clement yde, no (79, 83) illustrations: stefan sølberg, with contributions from jette halskov, eva melskens and benny m. schark layout and graphic production: annabeth andersen printers: rosendahls . schultz grafisk a/s, albertslund, denmark manuscripts submitted: 13 january – 16 april 2010 final versions approved: february–may 2010 printed: 7 july 2010 issn 1603-9769 (review of survey activities) issn 1604-8156 (geological survey of denmark and greenland bulletin) isbn 978-87-7871-282-0 citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 20, 106 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk and at www.geus.dk/publications/bull © de nationale geologiske undersøgelser for danmark og grønland (geus), 2010 for the full text of the geus copyright clause, please refer to www.geus.dk/publications/bull 3 44 ghana uganda tanzania mozambique greenland canada uk norway faroe islands kenya spain germany sw ed en the netherlands denmark brazil zambia belgium bolivia bahrain yemen sudan trinidad 7 review of survey activities 2009 f.g. christiansen 9 obituary: knud ellitsgaard-rasmussen, 23 june 1923 – 1 december 2009 n. henriksen and t.c.r. pulvertaft 15 late cretaceous basin development of the southern danish central graben f. jakobsen and c. andersen 19 identifying potential geothermal reservoirs in denmark a. mathiesen, l.h. nielsen and t. bidstrup 23 distribution and grain size of sand in the miocene wave-dominated billund delta, denmark e.s. rasmussen and j. bruun-petersen 27 3-d geological modelling of the egebjerg area, denmark, based on hydrogeophysical data f. jørgensen, r.r. møller, p.b.e. sandersen and l. nebel 31 late quaternary geology of a potential wind-farm area in the kattegat, southern scandinavia j.o. leth and b. novak 35 amino acid analysis of pre-holocene foraminifera from kriegers flak in the baltic sea o. bennike and b. wagner 39 radon content in danish till deposits: relationship with redox conditions and age p. gravesen and p.r. jakobsen 43 recent changes in the nutrient status of a soft-water lobelia lake, hampen sø, denmark k. weckström, p. rasmussen, b.v. odgaard, t.j. andersen, t. virtanen and j. olsen 47 silica diagenesis and its effect on porosity of upper maastrichtian chalk – an example from the eldfisk field, the north sea h.b. madsen 51 the continental shelf project of the kingdom of denmark – status at the beginning of 2010 c. marcussen and m.v. heinesen 55 greenland ice sheet monitoring network (glisn): a seismological approach t. dahl-jensen, t.b. larsen, p.h. voss and the glisn group 5 thailand vietnam phillipines cambodia malaysia sri lanka laos pdr china 59 the mineral resource assessment project, south-east greenland: year one b.m. stensgaard, j. kolb, t.f.d. nielsen, s.d. olsen, l. pilbeam, d. lieber and a. clausen 63 characterisation of host rocks and hydrothermal alteration of the qussuk gold occurrence, southern west greenland d.m. schlatter and r. christensen 67 zircon record of the igneous and metamorphic history of the fiskenæsset anorthosite complex in southern west greenland n. keulen, t. næraa, t.f. kokfelt, j.c. schumacher and a. scherstén 71 application of airborne hyperspectral data to mineral exploration in north-east greenland t. tukiainen and b. thomassen 75 study of a palaeogene intrabasaltic sedimentary unit in southern east greenland: from 3-d photogeology to micropetrography h. vosgerau, p. guarnieri, r. weibel, m. larsen, c. dennehy, e.v. sørensen and c. knudsen 79 an advancing glacier in a recessive ice regime: berlingske bræ, north-west greenland p.r. dawes and d. van as 83 bathymetry, shallow seismic profiling and sediment coring in sermilik near helheimgletscher, south-east greenland c.s. andresen, n. nørgaard-pedersen, j.b. jensen and b. larsen 87 borax – an alternative to mercury for gold extraction by small-scale miners: introducing the method in tanzania p.w.u. appel and j.b. jønsson 91 vietnamese sedimentary basins: geological evolution and petroleum potential m.b.w. fyhn, h.i. petersen, a. mathiesen, l.h. nielsen, s.a.s. pedersen, s. lindström, j.a. bojesen-koefoed, i. abatzis and l.o. boldreel 95 potential for permanent geological storage of co2 in china: the coach project n.e. poulsen 99 thin-skinned thrust-fault tectonics offshore south-west vietnam s.a.s. pedersen, l.o. boldreel, e.b. madsen, m.b. filtenborg and l.h. nielsen interactive web analysis and presentation of computer-controlled scanning electron microscopy data p. riisager, n. keulen, u. larsen, r.k. mclimans, c. knudsen and j. tulstrup geus working areas 2009. orange areas are covered in this volume. for further information on other working areas please refer to www.geus.dk/international 103 66 7 2009 was a favourable year for the geological survey of denmark and greenland (geus) with focus on research, often in international collaboration. many new projects have been initiated and many completed. in 2009, copenhagen hosted cop15, and geus’ involvement in the preparation for this event focussed on climate changes, reducing consumption of fossil fuels and co2 emissions. this is also reflected in this seventh issue of review of survey activities that describes many projects that geus and its partners carry out in denmark, greenland and abroad. together with the previous six issues, it provides a good overview of the survey’s range of research and advisory activities. the review contains a total of 23 four-page papers: nine on scandinavia, primarily denmark, one on the continental shelf project, eight on greenland, four on international work and one on data management. in addition, an obituary about knud ellitsgaard-rasmussen, 1923–2009, is given. he was the first leader and subsequently formal director of the geological survey of greenland (ggu) until he retired in 1983. the obituary gives an interesting overview of the post-war initiation and early history of ggu – merged in 1995 with the geological survey of denmark (dgu) to form the present geological survey of denmark and greenland. energy policy is again high on the political agenda in denmark. the financial crisis has emphasised the importance of revenues from the north sea oil production, at a time when we also see strong national and international political demands for reduction of co2 emissions. geus gives high priority to research within both topics – but also to research on climate development, climate monitoring and adaptation to climate changes. two papers concentrate on petroleum geology in denmark. one provides an overview of the late cretaceous basin development of the southern danish central graben where most of the danish oil and gas producing fields are found; the other discusses silica diagenesis and its effect on the porosity of impure maastrichtian chalk. denmark has a large potential for subsurface geothermal energy. geus and the danish energy agency have conducted a regional study to update the assessment of this potential; this work is summarised here, discussing future possibilities, including the geological risks and uncertainties. wind power is another important energy source that contributes to co2 reduction. most future wind power will probably come from large offshore wind farms. geus has been involved in the preparation of several projects, including evaluation of seabed features and subsurface characteristics of possible sites. one paper describes a case story from the kattegat. geus works on many other aspects of the geology of denmark, in particular in relation to groundwater and the environment. one paper describes the distribution and grain size of sand in the miocene billund delta. understanding these depositional systems is scientifically important and applicable for modelling of groundwater and petroleum resources. a second paper presents a detailed geological model of an area in eastern jylland, demonstrating how skytem, stratigraphic and lithological data have been integrated to provide better input to groundwater models. a third paper is on pre-holocene sediments from the baltic sea, discussing amino acid analysis of foraminifera and their application to age determination of the deposits. a fourth paper presents new results on the radon content in till deposits that seems to be controlled by age and redox conditions. this is important when evaluating health risks in houses in denmark. finally, the timing and possible causes of past nutrient enrichment in a lake in jylland have been investigated. such studies are essential to understanding the environmental threat to the ecosystem of the lake. 2009 was a historic year for the continental shelf project of the kingdom of denmark. the first partial claim (the area north of the faroe islands) was submitted to the commission of the limit of the continental shelf (clcs) in the spring. furthermore, a large number of data-acquisition surveys were completed, including lomrog ii with the swedish icebreaker oden that visited the north pole area. review of survey activities 2009 flemming g. christiansen deputy director © geus, 2010. geological survey of denmark and greenland bulletin 20, 7–8. open access: www.geus.dk/publications/bull 88 data acquisition is now complete south of the faroe islands and south of greenland. the technical work for the next submissions is well underway, but additional data from north-east and north of greenland are still required. a status paper is given on the project, which started in 2003 and has progressed as planned. the project will continue until 2014, possibly requiring follow-up work and data management for several years until clcs has submitted its recommendations to the united nations. in 2009, there was a high level of field activities onand offshore greenland. in addition to large campaigns in southern west, south-east, and east greenland there were many other smaller activities. results of ongoing work in southern west greenland are described in two papers, one on the zircon record of the igneous and metamorphic history of the fiskenæsset anorthosite complex, the other on the hydrothermal alteration and host rocks of a gold occurrence. 2009 was the first year in a mineral resource assessment project that is carried out together with the greenland bureau of minerals and petroleum in one of the least known regions of greenland, namely the logistically challenging south-east coast. a paper describes the start of the reconnaissance work that focussed on collecting stream sediments for an assessment of the economic potential of the area. field work and shallow core drilling in north-east greenland continued in 2009. related ore geology studies are described in a paper on application of airborne hyperspectral data in mineral exploration. in another part of east greenland, geus has been involved in 3-d photogeological studies of analogues to intrabasaltic reservoir sandstones elsewhere in the north atlantic region; these studies are presented in a paper that gives a good 3-d understanding and also provides details on reservoir properties. over many years, studies of the ice sheet and glaciers in greenland have attracted strong international interest due to the implications of global sea-level rise. geus is involved in many glaciological and meteorological projects and monitoring programmes. three papers focus on ice and climate. the first describes the greenland ice sheet monitoring network (glisn) that records glacial earthquakes and provides important information on changes in glacier dynamics. the second presents results on bathymetry, shallow profiling and sediment coring in sermilik near helheimgletscher in southeast greenland; work essential for studies of past ice fluctuations and their causes. the third is on glaciers in north-west greenland with a case story on the advancing berlingske bræ glacier in a generally recessive ice regime. geus works in many countries and with many projects types and has been active in vietnam for several decades, in later years especially in the so-called enreca project that has completed two out of its three planned phases. a paper describes vietnamese sedimentary basins, their geological evolution and petroleum potential. another deals with offshore, thin-skinned thrust-fault tectonics. small-scale mining provides income for millions of people in africa and asia, but often with negative effects on the environment and health of miners, especially from the widespread use of mercury in gold extraction. geus has been involved in projects with the overall goal of reducing, or even stopping, the release of mercury into the environment. a case story from tanzania with an alternative borax method is presented. the survey has also been involved in many euand industry-financed projects of carbon-dioxide capture and storage in the past decade. the coach project aims at applying this knowledge in china with its rapidly growing energy demand and where the use of coal-powered power stations is likely to continue for many years. a paper gives a first overview of co2 sources, proposed pipe lines and potential storage sites in eastern china. the final paper describes the results of a project carried out in cooperation with dupont titanium technologies featuring interactive web applications of scanning electron microscopy data. this project has been successful, and similar methods are likely to be used in mineral and petroleum exploration and for other projects with large databases. geological survey of denmark and greenland bulletin 4, 2003, pp 77-80 the discovery in 2002 of a gold mineralised quartz-carbonate vein at ubekendt ejland, central west greenland, yielding 0.6 ppm au over 0.7 m, led to a reconnaissance sampling project in summer 2003. most of the accessible quartz-carbonate veins on the south-east coast of the island (figs 1, 2) were sampled during boat-supported field work. massive sulphide mineral deposits (fe-zn-pb) were located in the centre of brecciated quartz-carbonate vein systems at several places along the south and south-east coast of the island, and gold anomalies mainly associated with the occurrence of the massive sulphides were identified. pervasive hydrothermal alteration of the volcanic wall rocks surrounds the quartz-carbonate vein systems, which comprise low-temperature mineral assemblages dominated by dolomite and veined by chalcedony and fibrous silica. evidence of oil migration into volcaniclastic rocks prior to the intense hydrothermal activity was found in several places in the form of organic carbon, interpreted to be pyrobitumen, that infills pores and cavities in hyaloclastites. geological setting ubekendt ejland comprises early palaeogene volcanic and intrusive rocks (fig. 1; drever & game 1948; larsen 1977a, b). 77 epithermal gold and massive sulphide mineralisation in oil impregnated palaeogene volcanic rocks of ubekendt ejland,west greenland stefan bernstein and christian knudsen geological survey of denmark and greenland bulletin 4, 77–80 (2004) © geus, 2004 fig. 1. geological map of ubekendt ejland. most quartz-carbonate veins (red lines) visited in 2003 are located from 5 km east of nuunngutak to 7 km west of qeqertalik. modified from larsen (1977a, b). spot heights (e.g. 420) are in metres. most of the island consists of picritic and olivine-phyric lavas, hyaloclastites and volcanic breccias, assigned to the vaigat formation (fig. 1). the vaigat formation succession dips 20–30° to the west and is overlain by plagioclase-phyric lavas, pyroclastic units and an upper lava succession including alkaline basalts (larsen 1977b). a mafic–felsic intrusive complex, that includes layered gabbros and fine-grained granite, is found at saqqaata qaqqaa in the southern part of the island. the quartz-carbonate veins are particularly abundant in the south-eastern part of the island, but veins also occur close to the south-west coast, and on the north-west coast at qarusuk (fig. 1). most veins on the south-eastern coast can be traced for a few kilometres northwards until they are hidden beneath quaternary glacial deposits. one vein system has been traced for 6 km along strike. the highest density of vein systems occurs along a 5 km long stretch of the coast east of nuunngutak. quartz-carbonate veins some vein systems appear to be developed as relatively welldefined planar structures that extend for some distance along strike, judging from the orange-buff outcrop coloration. in other cases, the veins form complex anastomosing systems, where sets of veins – often four or more – follow irregular trends and join upwards (fig. 2). in general, the vein systems are subvertical. it is unclear whether the vein systems were formed during one or several events, although at several locations the veins exhibit evidence of distinct stages of brecciation and mineralisation (see below). samples were taken up to about 5 km west of the intrusive complex at saqqaata qaqqaa, but no significant changes in mineralisation assemblages or structures were observed. the quartz-carbonate veins on ubekendt ejland typically consist of banded carbonates with centres filled by quartz or carbonate. thin (< 1 mm) cross-cutting veins filled with quartz and/or fibrous silica also occur. the banded carbonates are mediumto fine-grained and some veins also have coarsegrained centres with carbonate crystals up to 20 mm across (fig. 3). the outer margins of the veins are often intensely brecciated, showing a mixture of carbonate and quartz crystals and fragments, including fragments of chalcedony, set in a very fine-grained matrix of ground carbonate. the matrix sometimes shows signs of recrystallisation. carbonate is predominantly dolomite, and grades into ankerite. some quartzcarbonate veins exhibit later sulphide mineralisation (see below) lining late fractures and veins (fig. 3). the late fracturing can be extensive and has resulted in brecciation zones within the quartz-carbonate veins. the latest veining and mineralisation stage resulted in the formation of chalcedony veins that are usually 1–2 mm thick, but chalcedony commonly occurs in cavities and vein centres within massive sulphides. in places, the chalcedony grades into coarser-grained quartz (0.5 mm). vugs filled with fibrous quartz, probably replacing amorphous sillica, are also common. sulphide and gold mineralisation in the larger quartz-carbonate veins, the sulphide mineralised fractures and breccias may be up to 15 cm in width and 40 cm long and comprise semi-massive to massive sulphides. in brecciated quartz-carbonate veins, pyrite and other sulphide minerals are euhedral or subhedral and fine-grained, while in 78 chalcedony sulphides fig. 2. quartz-carbonate vein systems cutting picritic lava flows of the vaigat formation, about 3 km east of nuunngutak (fig. 1). height of cliff about 150 m. the orange-buff coloured patterns include both the quartzcarbonate veins and their alteration halos, typically 2–10 m wide. fig. 3. a 20 cm thick quartz-carbonate vein in an altered dyke, 6 km west of qeqertalik (fig. 1). the vein shows several stages of brecciation, with fragments of banded dolomite overgrown with pyrite. some cavities are lined with chalcedony, while others are lined with coarse calcite crystals. pen is 15 cm long. the massive sulphides, pyrite, pyrrhotite and sphalerite are subhedral and coarse-grained, most often enclosing other sulphides such as galena, arsenopyrite, millerite, pentlandite and chalcopyrite. in most vein systems, the sulphides occur as veins and in the brecciated matrix within the quartz-carbonate veins, but disseminated sulphides also occur in the altered wall rocks. fresh sulphides have only been found in coastal outcrops, while elsewhere they are extensively altered to limonite. massive sulphide bodies comprise mainly pyrrhotite or pyrite, sphalerite, galena, arsenopyrite and native silver. gold has not been seen in thin section, but samples of the massive sulphides contain up to 1300 ppb au and 110 ppm ag, with 3.8 wt% pb and 2.2 wt% zn. the gold appears to reside within the sulphide-mineralised vein centres. alteration of wall rocks the igneous host rocks are lava flows, volcaniclastites, hyaloclastites and basaltic dykes, all of which have suffered extensive hydrothermal alteration adjacent to the quartz-carbonate veins. the alteration halos may extend up to tens of metres, but more usually are in the range of 2–5 m on either side of the vein system. veins that follow older thick basaltic dykes often only show hydrothermal alteration in a narrow (1–2 m) zone adjacent to the vein, while the remaining dyke retains its igneous mineral assemblage. at one location, a hyaloclastite unit is cut by a vertical, 2 m thick basaltic dyke. the eastern dyke contact is cut by a parallel quartz-carbonate vein system, and alteration extends to the centre of the dyke. on the east side of the vein, alteration within the hyaloclastite host rock extends for about 8–10 m. the unaltered part of the basaltic dyke contains sparse, fresh clinopyroxene phenocrysts set in a groundmass of finegrained plagioclase, clinopyroxene and fe-ti oxide grains. intersticies are filled with sericite, and vesicles with chalcedony, carbonate and illite/phengite. in the altered part of the dyke, plagioclase laths in the groundmass are still visible, but are invariably altered to sericite. former clinopyroxene phenocrysts are replaced by fine-grained yellow clay minerals, and bands of fine-grained euhedral pyrite cross-cut the rock, together with thin (50 micron) veins of carbonate or chalcedony. chemical analyses demonstrate that the alteration has resulted in the loss of sio2, mgo, cao, na2o, sr and ba, and in an increase in feototal, k2o, ni, rb, pb and volatiles. other elements appear unaffected by the alteration; these include ti, al, p, the transition metals, zr and the rare earth elements. these changes reflect the dissolution of igneous silicates and growth of hydrous phyllosilicates, carbonate and pyrite. we deduce that the added volatiles must include h2o, co2 and s. trace of oil on ubekendt ejland evidence of oil migration into the onshore palaeogene volcanic rocks is common in central west greenland (bojesenkoefoed et al. 1999). on the east coast of ubekendt ejland (fig. 1), minor oil staining has earlier been reported at the contact of a dyke (christiansen et al. 1998). during field work in 2003, several localities with quartz-carbonate veins and alteration zones on the south-east coast of ubekendt ejland were found to bear evidence of oil migration into the 79 fig. 4. upper portion of altered hyaloclastite, some 3 m from a quartz-carbonate vein. the black matrix between the lithic fragments is organic carbon, which also fills fine cracks in the host rock. to the right is seen a subset of thin carbonate veins. pencil is 15 cm long. about 4 km east of nuunngutak. 80 volcanic rocks. at one location, this is manifested by the patchy development of black staining in altered hyaloclastic rocks (fig. 4). in thin section, organic carbon occurs as almost opaque material enveloping lithic fragments identical to the surrounding altered hyaloclastite (fig. 5). analysis of the rock material yielded 1.25 wt% organic carbon, 2.89 wt% carbon and 0.14 wt% sulphur. these patchy zones in the hyaloclastites are interpreted to be remnants of oil (pyrobitumen) that migrated into the volcanic rocks and filled pore space in the hyaloclastites. in this case the migration of oil occurred some time before the intense hydrothermal alteration associated with the emplacement of the quartz-carbonate vein systems. prolonged exposure to elevated temperatures has caused thermal alteration of the oil to leave pyrobitumen as the solid residue. assuming a density of the pyrobitumen of about 1 g/cm3, the 1.25 wt% carbon corresponds to about 3 vol.% organic carbon in the rock, which in turn would represent 10% of the original volume of oil. an inferred 30 vol.% porosity in the matrix of the hyaloclastites appears likely, either reflecting primary porosity or attained by dissolution of primary pore fillings prior to the invasion of oil. conclusion the presence of chalcedony and fibrous silica possibly replacing amorphous silica suggest that the quartz-carbonate veins formed at a shallow level in hydrothermal conduits. the quartz-carbonate veins are similar to epithermal au-rich vein mineralisations described from other volcanic terrains associated with organic-rich sediments, in terms of geological setting, vein mineralogy, alteration of wall rocks and the evidence of migration of hydrocarbons (see e.g. sketchley & sinclair 1991; sherlock & lehrman 1995; sillitoe et al. 2002). the vein mineralogy and alteration assemblage further suggest the mineralisation is a low sulphidation type (e.g. bonham 1988). the presence of fe-pb-zn sulphides in several veins may indicate that the sampled level of the veins is below that of precious metal deposition, that according to the model of buchanan (1981) would appear above that of base metal sulphides. our findings suggest there is a potential for economic precious metal mineralisation on ubekendt ejland. references bojesen-koefoed, j.a., christiansen, f.g., nytoft, h.p. & pedersen, a.k. 1999: oil seepage onshore west greenland: evidence of multiple source rocks and oil mixing. in: fleet, a.j. & boldy, s.a.r. (eds): petroleum geology of northwest europe: proceedings of the 5th conference, 305–314. london: geological society. bonham, h.f. 1988: models for volcanic-hosted epithermal precious metal deposits. in: schafer, r.w., cooper, j.j. & wikre, p.g. (eds): bulk mineable precious metal deposits of the western united states, 259–272. reno: geological society of nevada. buchanan, l. 1981: precious metal deposits associated with volcanic environments in the southwest. ar geological society digest 14, 237–262. christiansen, f.g., boesen, a., bojesen-koefoed, j.a., dalhoff, f., dam, g., neuhoff, p.s., pedersen, a.k., pedersen, g.k., stannius, l.s. & zinck-jørgensen, k. 1998: petroleum geological activities onshore west greenland in 1997. geology of greenland survey bulletin 180, 10–17. drever, h.i. & game, p.m. 1948: the geology of ubekendt ejland, west greenland. meddelelser om grønland 134(8), 34 pp. larsen, j.g. 1977a: fieldwork on ubekendt ejland in the tertiary basalt province of west greenland, 1971 and 1973. rapport grønlands geologiske undersøgelse 79, 35–44. larsen, j.g. 1977b: transition from low potassium olivine tholeiites to alkali basalts on ubekendt ejland. meddelelser om grønland 200(1), 42 pp. sherlock, r.l. & lehrman, n.l. 1995: occurrences of dendritic gold at the mclaughlin mine hot-spring gold deposit. mineralium deposita 30, 323–327. sillitoe, r.h., cooper, c., sale, m.j., soechting, w., echavarria, d. & gallardo, j.l. 2002: discovery and geology of the esquel low-sulphidation epithermal gold deposit, patagonia, argentina. society of economic geologists, special publication 9, 227–240. sketchley, d.a. & sinclair, a.j. 1991: carbonate alteration in basalt, total erickson gold mine, cassiar, northern british columbia, canada. economic geology 86, 570–587. fig. 5. thin section of sample from locality in fig. 4 (plane-polarised light). the dark matrix contains lithic fragments of varying size and of a similar composition to the host rock. note the fine veins (arrowed) of dark carbon-rich material extending into the altered hyaloclastite. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: sb@geus.dk << /ascii85encodepages false 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<feff4f7f752890194e9b8a2d5b9a5efa7acb76840020005000440046002065874ef65305542b8f039ad876845f7150cf89e367905ea6ff0c4fbf65bc63d066075217537054c18cea3002005000440046002065874ef653ef4ee54f7f75280020004100630072006f0062006100740020548c002000520065006100640065007200200035002e0030002053ca66f465b07248672c4f86958b555f3002> >> >> setdistillerparams << /hwresolution [2400 2400] /pagesize [595.000 842.000] >> setpagedevice geological survey of denmark and greenland bulletin 7, 2004, p 53-56 53 in the present-day western north atlantic, icebergs can be observed off north-east canada, drifting south along the coast in the cold labrador current. normally they melt in the area off newfoundland where they reach warmer waters. most of these icebergs originate from calving glaciers in west greenland or in the canadian arctic. jakobshavn isbræ in west greenland (fig. 1) deserves particular mention as it is the fastest known ice stream in the world draining 6–7% of the entire greenland ice sheet (joughin et al. 2004). southward drifting icebergs also occur along the east coast of greenland (fig. 2), but most of these melt when they approach the southernmost tip of greenland. the iceberg limit in the north-western atlantic varies from year to year, but isolated icebergs may reach far south of newfoundland (fig. 1). many icebergs carry a load of rock debris and soil incorporated by their parent glacier that leads to deposition of ice rafted debris on the deep ocean sea floor. in the past decade the geological survey of denmark and greenland (geus) has initiated marine geological investigations in the north atlantic on the late quaternary variability of north atlantic thermohaline circulation, with special focus on the possible link between climate change and variations in deep-water flow intensity (kuijpers et al. 1998, 2002, 2003). moreover, glaciological projects in greenland undertaken by geus have significantly contributed to the current debate of present-day climatic warming. notably work carried out in east greenland fjords has provided crucial information relevant for the study of glacial iceberg surges in the north atlantic (reeh et al. 1999). these surges are suggested to have been triggered under the influence of extreme cold climate conditions, but the actual trigger mechanism involved has been a matter of much debate. evidence from modern glacier process studies referred to above, combined with results of recent studies in the north atlantic carried out by geus and partner institutions, has provided new insights into the possible trigger mechanism of these massive glacial iceberg surges. these new findings have great significance for the current climate debate, since they strongly suggest that ongoing ocean warming can trigger a sudden, massive break-up of ice shelves. such processes may already be in progress in the arctic (e.g. vincent et al. 2004), where rapid ice-shelf disruption on the margin of the caclimatic warming: a trigger for glacial iceberg surges (‘heinrich events’) in the north atlantic? antoon kuijpers, hartmut heinrich and matthias moros geological survey of denmark and greenland bulletin 7, 53–56 (2005) © geus, 2005 greenland newfoundland baffin island canada 40°w 40°n 80°n 0° norway spain iceland faroe islands great britain jakobshavn isbræ hudson strait labrador sea east gree nla nd curr ent labrador current iso w n ac bay of biscay greenland sea 1881 1881 1936 1903 1917 1944 1912 1901 1912 r.m.s. titanic 1913 1918 1921 1948 1908 1886 1842 1907 1 2 3 fig. 1. overview of the north atlantic with iceberg drift pattern and exceptional iceberg observations. the main deep-water flow path of iceland–scotland overflow water (isow) and the study areas (1–3) mentioned in the text are also indicated. locality 1 refers to the area north-west of spain studied by heinrich (1988), whereas 2 and 3 indicate study areas located along the isow deep-water flow path south of iceland. nac (arrow) shows northward transport of warm, saline water in the upper water masses by the north atlantic current. pale blue indicates areas dominated by seasonal drift of icebergs. nadian arctic ocean has been reported to be the result of significant warming over the past few decades. during this period intensified inflow of atlantic water to the eurasian sector of the arctic has been noted. it is evident that for antarctic ice shelves large-scale disruption and break-up may lead to significant destabilisation of the antarctic ice sheet with the serious risk of a sudden, drastic sea-level rise. ice rafted debris layers and ‘heinrich events’ in the beginning of the 1980s several west european governments spent considerable funds on marine (geo)science investigating the possible option of future sub-seabed disposal of high-level radioactive waste, and the environmental risks of dumping low radioactive waste on the north-east atlantic seabed. the german government supported marine geological investigations in an area of the eastern north atlantic which had similar characteristics to the official nuclear energy agency dump sites in the bay of biscay (fig. 1). up to 22 m long sediment cores were retrieved from water depths ranging from 3500 to 4500 m, and in the upper part of the cored sediment sequence a series of distinct layers rich in ice rafted debris (ird) were identified (fig. 3). the ird layers mainly consist of sand and were formed during the last glacial cycle (weichselian), which started around 110 000 years ago. these layers were further characterised by containing only very few tests of polar planktonic microfauna (foraminifera), much less than in the surrounding glacial sediments (heinrich 1988). the estimated ages of the upper six ird layers roughly correspond to times of major changes in the sea surface temperature of the northern atlantic during the last glacial period. the lowermost ird layer marks the initiation of full glacial conditions, and the uppermost layer the termination of the last ice age (heinrich 1988). soon after the publication of these findings similar layers were reported from sediment cores taken in the labrador sea (bond et al. 1992; broecker 1994). subsequent investigations showed that the sand-size fractions of these layers were also characterised by elevated contents of detrital carbonate. mineralogical investigations of the fine fraction of these deposits indicate that the ird of some of these layers originated from the greenland and north american region, while two of the layers included compounds of volcanic weathering products, possibly from the european arctic (jantschik & huon 1992). mineral ages (k-ar) of the ird were much higher than measured for surrounding glacial sediments. another important observation was that the ird layers thin by more than an order of magnitude from the labrador sea towards the eastern north atlantic (grousset et al. 1993), and that north atlantic surface waters were extremely cold during their deposition. the latter conclusion was based on the high percentage of the polar planktonic foraminifera neogloboquadrina pachyderma (sinistral), and is supported by low concentrations of 18o in the foraminiferal calcium carbonate tests (fig. 3). the drastic lowering of about 5°c of surface water temperature as well as the decrease of surface water salinity were thought to result from the melting of icebergs originating mainly from massive calving of the north american (laurentide) ice sheet around hudson strait. formation and south-eastward expansion of a cold meltwater lid may have led to the weakening or cessation of the north atlantic thermohaline circulation system. thus, transport of warm (saline) surface water via the north atlantic current (nac) to northern high latitudes decreased, resulting in a dramatic cooling in europe. a detailed chronology established for these recurrent ird events shows a spacing of about 11 000 years during the early glacial and about 7000 years during full glacial conditions. the significance of these massive iceberg discharge events for the understanding of the global climate evolution of the last glacial era led to the introduction of the term ‘heinrich events’ (h-events) for these extreme glacial iceberg episodes in the north atlantic (broecker et al. 1992). in addition, possible links between h-events in the north atlantic and the so-called dansgaard-oeschger (d-o) cycles recorded in the greenland ice cores were investigated by bond et al. (1993). the glacial greenland ice core record (dansgaard et al. 1993) shows marked millennial scale (1500 years) climatic cycles, revealing abrupt warming (5–10°c) within a few decades, subsequently followed by slow cooling and terminating in an extremely cold (stadial) period lasting tens to hundreds of 54 fig. 2. the margin of the inland ice in south-east greenland with outlet glaciers and numerous icebergs. widespread slush (white) is observed in areas close to the coast. years. bond et al. (1993) demonstrated that h-events coincide with the coldest parts of a set of several d-o cycles and were immediately followed by significant warming. meanwhile, large-scale glacial iceberg surges have been found not to be exclusively related to the north american region, but have also been documented for the north-west european ice sheet (e.g. knutz et al. 2002). ocean subsurface warming and ice-shelf break-up during the past decade numerous studies have been published dealing with possible trigger mechanisms for h-events. possible explanations range from internal ice-sheet mechanisms (binge purge theory) to external forcing processes such as ice-load induced earthquakes, sea-level change, variations in solar forcing, enhanced tides and other mechanisms summarised by alley & clark (1999). moros et al. (2002) proposed an alternative, new theory – outlined below – that relies on the observations of recent iceberg discharge processes and ice-rafting in east greenland fjords reported by reeh et al. (1999), who also described the implications of advection of atlantic intermediate water at subsurface depth (> 200 m) for glacier and iceberg bottom melting and ird distribution. similarly, changes in glacial ocean subsurface circulation and associated heat transport play a key role in the theory proposed by moros et al. (2002). as part of the global ocean thermohaline circulation system the north atlantic current (nac) transports heat and salt via an upper flow (< c. 800 m) into northern high latitudes. after cooling in the labrador sea and greenland sea the density of the upper water masses increases, leading to convection and formation of cold, dense water flowing south as deep countercurrents. the occurrence of high salinities in the greenland sea is a crucial factor driving deep-water convection; low surface water salinity will lead to a weakening of the formation of deep, cold water. on their way south, the dense and cold water masses formed in the greenland sea can be traced in the northern north atlantic as iceland–scotland overflow water (isow). by studying sediment records from cores taken along the isow flow path (fig. 1, localities 2 and 3), moros et al. (2002) demonstrated that isow flow intensified immediately before an h-event (fig. 4). in deep-sea sediments the silt fraction and other lithological parameters have proven to be a sensitive proxy for changing bottom water dynamics. moros et al. (2002) found that increased flow activity, indicated by a coarsening of the silt fraction, precedes the maximum (continental) ird input reflected by the quartz/plagioclase peak (fig. 4). thus, isow bottom current activity increased just before large-scale iceberg surging. due to the coupling between the north atlantic current and isow, an increase in isow indicates that nac also increased. intensification of both deep isow and high nac flow is a typical feature of climate warming (e.g. kuijpers et al. 1998). during cold glacial climate times, the northward flow of warmer and more saline atlantic water must initially have occurred below a cold surface layer, similar to that observed in the arctic ocean at present. this enhanced northward subsurface heat transport would increase bottom melting of floating outlet glaciers and ice shelves, ultimately resulting in ice-sheet destabilisation and massive iceberg surging. for icebergs to survive over large distances, as reported by the widespread ird layers, initial ice melting must have been fast, creating an extensive cold, lowsalinity surface layer. thus, a marked cooling of the surface water was the consequence of initial, large-scale subsurface warming and melting. at lower latitudes in the north atlantic simultaneous warming has been reported by various authors. this could explain the sudden rise in temperature a short time after the h-events, when the iceberg supply came to an end and the low salinity lid became fully mixed with the warmer water. the same mechanism of subsurface warming 55 0 20 40 60 % 80 100 0 0 1 2 3 420 40 60 % % 80 100 0 50 100 150 [c m ] 200 250 300 0 50 100 150 200 250 300 0 50 100 150 200 250 300 g. truncatulinoides g. inflata n. pachyderma (l.) ∂18 o g. bulloidesice-rafted debris he 1 he 2 he 3 he 4 he 5 he 6 8 10 11 me69-17 fig. 3. core record me69-17 from the area north-west of spain (locality 1 in fig. 1). the amount of ice-rafted debris with peaks representing heinrich events 1 to 6 (he 1–6) is shown as part of the total split (ird plus foraminifera), whereas foraminiferal species abundance is presented as part of the sum of the planktonic foraminiferal fauna. in addition, the stable oxygen isotope record of g. bulloides is shown. slightly modified from heinrich (1988), reproduced with permission from elsevier. 56 as a trigger for outlet glacier melting and ird production has recently been documented for the early holocene retreat of jakobshavn isbræ (lloyd et al. in press). future work as outlined above, geus studies have hitherto been strongly focused on the northern north atlantic. however, studies that deal with the reconstruction of oceanographic processes and iceberg drift in the labrador sea and baffin bay region have recently been initiated, and will include multi-proxy analyses of sediment cores. acknowledgements m.m. acknowledges financial support from the german research council (dfg) and the european marie curie fellowship funding agency. references alley, r.b. & clark, p.u. 1999: the deglaciation of the northern hemisphere: a global perspective. annual review earth and planetary sciences 27, 149–182. bond, g.c. et al. 1992: evidence of massive discharges of icebergs into the north atlantic ocean during the last glacial period. nature 360, 245-249. bond, g.c., broecker, w.s., johnsen, s., mcmanus, j., labeyrie, l., jouzel, j. & bonani, g. 1993: correlations between climate records from north atlantic sediments and greenland ice. nature 365, 143–147. broecker, w.s. 1994: massive iceberg discharges as triggers for global climate change. nature 372, 421–424. broecker, w.s., bond, g., klas, m., clark, e. & mcmanus, j. 1992: origin of the northern atlantic’s heinrich events. climate dynamics 6, 265273. dansgaard, w. et al. 1993: evidence for instability of past climate from a 250 kyr ice core. nature 364, 218–220. grousset, f., labeyrie, l., sinko, j.a., cremer, m., bond, g., duprat, j., cortijo, e. & huon, s. 1993: patterns of ice rafted detritus in the glacial north atlantic (40°-50°n). paleoceanography 8, 175-192. heinrich, h. 1988: origin and consequences of cyclic ice rafting in the northeast atlantic ocean during the past 130,000 years. quaternary research 29, 142–152. jantschik, r. & huon, s. 1992: detrital silicates in northeast atlantic deep sea sediments during the late quaternary: mineralogical and k-ar isotope data. eclogae geologica helvetica 85, 195-212. joughin, i., abdalati, w. & fahnstock, m. 2004: large fluctuations in speed on greenland’s jakobshavn isbræ glacier. nature 432, 606-610. knutz, p., hall, i.r., zahn, r., rasmussen, t.l., kuijpers, a., moros, m. & shackleton, n.j. 2002: multidecadal ocean variability and nw european ice sheet surges during the last deglaciation. geochemistry, geophysics, geosystems 3(12), 1077 (http://dx.doi.org/10.1029/2002gc000351). kuijpers, a., troelstra, s.r., wisse, m., heier nielsen, s. & van weering, t.c.e. 1998: norwegian sea overflow variability and ne atlantic surface hydrography during the past 150,000 years. marine geology 152, 75–99. kuijpers, a., hansen, b., hühnerbach, v., larsen, b., nielsen, t. & werner, f. 2002: norwegian sea overflow through the faroe–shetland gateway as documented by its bedforms. marine geology 188, 147–164. kuijpers, a., troelstra, s.r., prins, m.a., linthout, k., akhmetzhanov, a., bouryak, s., bachmann, m.f., lassen, s., rasmussen, s. & jensen, j.b. 2003: late quaternary sedimentary processes and ocean circulation changes at the southeastern greenland margin. marine geology 195, 109–129. lloyd, j.m., park, l.a., kuijpers, a. & moros, m. in press: early holocene palaeoceanography and deglacial chronology of disko bugt, west greenland. quaternary science reviews 24. moros, m., kuijpers, a., snowball, i., lassen, s., bäckström, d., gingele, f. & mcmanus, j. 2002: were glacial iceberg surges in the north atlantic triggered by climatic warming? marine geology 192, 393–417. moros, m., mcmanus, j., rasmussen, t., kuijpers, a., dokken, t., snowball, i., nielsen, t. & jansen, e. 2004: quartz content and the quartz to plagioclase ratio determined by x-ray diffraction: a proxy for iceberg and sea ice rafting in the northern north atlantic. earth and planetary science letters 218, 389–401. reeh, n., mayer, c., miller, h., thomsen, h.h. & weidick, a. 1999: present and past climate control on fjord glaciations in greenland: implications for ird deposition in the sea. geophysical research letters 26, 1039-1042. vincent, w.f., van hove, p. & mueller, d.r. 2004: sensitivity to climate change in the canadian high arctic: ellesmere island lakes, fiords and ice shelf ecosystems. acia international scientific symposium on climate change in the arctic, reykjavik, iceland, 9–12 november, 2004. exended abstracts. amap report 4, session 5, paper 8. authors’ addresses a.k., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: aku@geus.dk h.h., federal maritime and hydrographic agency (bsh), bernhard-nocht-str. 78, d-20359 hamburg, germany. m.m., baltic sea research institute, seestrasse 15, d-18119 rostock, germany. h1 0 50 100 150 50 100 150 200 50 100 150 20050 100 150 silt grain size (q /p lag ratio ) iso w in te n sity iir d in p u t h1 h1h1 enam-30lo09-02gc so82-05ggc depth [cm] lo09-18gc yd ydyd fig. 4. sections from three core records from the study area south of iceland (lo and so cores) and from south-west of the faroe islands (enam core; for location see fig. 1). all show coarsening within the silt fraction indicating intensified flow of isow immediately before the h1 event. this occurred at the initial termination of the last deglaciation, slightly more than 15 000 years ago. the quartz / plagioclase (q/plag) peak indicates maximum ird. data from moros et al. (2004). geological survey of denmark and greenland bulletin 42, 2018, 115-126 115 provenance of basinal sandstones in the upper jurassic hareelv formation, jameson land basin, east greenland mette olivarius, morten bjerager, nynke keulen, christian knudsen and thomas f. kokfelt zircon u–pb geochronology and heavy mineral ccsem analysis were used to interpret the provenance of oxfordian–volgian sandstones of the hareelv formation in east greenland. six samples were collected from the blokelv-1 core drilled in southern jameson land, and the zircon age distributions and heavy-mineral assemblages are quite uniform. the samples contain a wide spectrum of archaean to palaeozoic zircon ages with peak ages at 2.71, 2.49, 1.95, 1.65, 1.49, 1.37, 1.10 and 0.43 ga when combining all data. the heavy-mineral compositions show derivation from felsic source rocks, some of which were metamorphic. the results reveal that the sediment was derived from the caledonides, and it is plausible that some or all of the material has experienced several cycles of sedimentation. devonian and carboniferous sediments preserved north of the area have zircon age distributions that correspond to those from the hareelv formation, and such rocks may have been reworked into the jameson land basin. the provenance signature describes both the gravity-flow sandstones of the hareelv formation and the delta-edge sands that are inferred to have fed them. lithological and provenance contrasts between the sandstones of the sjællandselv member and those of the katedralen member indicate a shorter transport distance, source to sink, suggestive of proximal topographic rejuvenation in the volgian. keywords: east greenland, hareelv formation, upper jurassic, sediment provenance, zircon geochronology, heavymineral analysis ___________________________________________________________________________ geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350 copenhagen, denmark. e-mail: mol@geus.dk the fully cored borehole blokelv-1 was drilled in 2008 to a depth of 233.8 m in southern jameson land, east greenland (fig. 1) to provide reservoir and source-rock information on the mudstone-dominated upper jurassic hareelv formation that also includes numerous gravityflow sandstones (bjerager et al. 2018a, this volume). the east greenland caledonides, which form the basement for the mesozoic succession, comprise a complex series of archaean to palaeozoic lithologies (higgins & leslie 2008), and therefore the post-caledonian sedimentary successions preserved along the east and north-east greenland coast (fig. 1) are expected to contain a wide range of zircon age populations. detrital zircon ages from jurassic sediments in east greenland have previously only been reported from the lower jurassic succession in the jameson land basin where the elevated liverpool land high was the primary source area (slama et al. 2011). significantly different zircon age distributions are to be expected in the upper jurassic sediments in the jameson land basin since the liverpool land high was gradually onlapped and drowned during the middle jurassic (surlyk 2003). the present study employs zircon geochronology and heavy-mineral analysis to characterise the upper jurassic sandstones of the hareelv formation in the blokelv-1 core and to identify the source areas of these sandstones. © geus, 2018. geological survey of denmark and greenland bulletin 42, 115–126. available at: www.geus.dk/bulletin42 mailto:mol@geus.dk http://www.geus.dk/bulletin42 116116 geological setting the caledonian orogen in east greenland formed during the palaeozoic closure of the iapetus ocean and continental collision between baltica and laurentia (mckerrow et al. 2000). the caledonian foreland of laurentia (greenland) consists mainly of crystalline rocks that are presently exposed in tectonic windows along the rim of the greenland ice sheet (fig. 1; henriksen et al. 2008). the caledonian orogen is composed of several westward displaced thrust sheets that consist of archaean to silurian crystalline and sedimentary rocks (higgins et al. 2004). the crystalline complexes in the orogen were formed during archaean and palaeoproterozoic times (thrane 2002). the sediments of the krummedal supracrustal sequence were deposited in the late mesoproterozoic to early neoproterozoic and underwent high-grade metamorphism, migmatisation and partial melting in the early neoproterozoic (kalsbeek et al. 2000, 2008a; watt et al. 2000). a thick sedimentary succession was subsequently deposited in the eastern part of the area which consists of the neoproterozoic eleonore bay supergroup overlain by the upper neoproterozoic tillite group and the lower palaeozoic kong oscar fjord group (higgins et al. 2004). this succession and the underlying basement units were partially affected by caledonian metamorphism. marine sedimentation was terminated during the uplift caused by the caledonian crustal thickening, and lateto post-caledonian continental sediments were deposited during the devonian after the extensional collapse of the caledonian orogen (larsen et al. 2008). the caledonian orogeny caused renewed granite intrusion and migmatisation during the ordovician–silurian (kalsbeek et al. 2001). fig. 1. geological map of the caledonian belt in east and north-east greenland based on stemmerik et al. (1997), henriksen et al. (2008) and kalsbeek et al. (2008b). the sampled blokelv-1 borehole is located in southern jameson land. note that the sediments and metamorphic rocks are categorised by their depositional ages. palaeogene volcanic rocks devonian continental neoproterozoic and caledonian granite cambrian–silurian neoproterozoic (mainly eleonore bay supergroup) meso–neoproterozoic metamorphic rocks (including krummedal supracrustal sequence) archaean–palaeoproterozoic crystalline complex tectonic windows with crystalline rocks thrust fault/shear zone cretaceous jurassic triassic permian carboniferous hareelv & raukelv formations 70°n 74°n jameson land in land ice caledonides kong oscar fjord 100 km25°w liverpool land sco re sb y sund h g f h h e g d c b b–h f localities for zircon age data from literature (see fig. 7) blokelv-1 borehole mol prov 1 117 the post-caledonian sedimentary succession of carboniferous to palaeogene age is exposed along the coast of east and north-east greenland from 70° to 76°n (fig. 1; e.g. stemmerik et al. 1992) except where it is overlain by extensive palaeogene plateau basalts that were extruded in association with the north atlantic continental breakup (brooks 2011). thermal subsidence in the sedimentary basins was succeeded by rifting in the middle jurassic and a thick succession was deposited in the jameson land basin until the rifting in the region waned in the earliest cretaceous (surlyk 2003). the oxfordian olympen formation with its marine basinal mudstones and deltaic sandstones is the only part of the upper jurassic succession that is preserved in northern jameson land (fig. 2), where it is exposed on the highest mountain peaks (larsen & surlyk 2003). the orientation of planar cross-bedding in the delta-top deposits in the upper part of the formation (zeus member) shows that the delta prograded towards the south. in southern jameson land, a near-complete upper jurassic succession is exposed and consists of basinal mudstones interbedded with gravity-flow sandstones of the oxfordian–kimmeridgian katedralen member of the hareelv formation (surlyk 2003). these sandstones were formed by collapse of the drowned shelf-edge delta deposits of the olympen formation situated at the northern margin of the jameson land basin (figs 2, 3) (bjerager et al. 2018b, this volume). some of the gravity-flow sandstones of the katedralen member were remobilised and intraformationally injected after burial (surlyk et al. 2007). the coarser-grained gravity-flow sandstones of the volgian sjællandselv member of the hareelv formation c re ta ce ou s lo w er u pp er ju ra ss ic o xf or di an l l l l m u u u u sy st em se rie s te th ys (s ta nd ar d) bo re al su bs ta ge m ki m m er id gia n be rr ias ian ti th on ian ry az an ian vo lgi an age (ma) 140 150 160 145 155 stage s njameson land hareelv fm? hareelv fm hades mb olympen fm zeus mbzeus mb sa katedralen mb sjællandselv mb shallow marine sandstone shelf transition silty–sandy mudstones and heterolith deep marine sandstone deep shelf (basin) mudstone mass-flow sandstone remobilised/injected sandstone raukelv fmraukelv fm id 6 id 5 id 4 id 3 id 2 id 1 samples (id 1–6) source rockprograding unit mol prov 2 bl ok el v1 fig. 2. stratigraphic scheme of the upper jurassic succession in jameson land based on bjerager et al. (2018b, this volume). 118118 were subsequently deposited in the jameson land basin (fig. 3) in response to a regional sea-level fall; southwards progradation of the shelf-edge clinothems of the raukelv formation into the basin led to repeated shelf-edge collapse and the initiation of the gravity flows (surlyk 2003; surlyk & noe-nygaard 2005). the provenance of the jurassic sedimentary basins in east greenland based on sedimentological data has been interpreted as the basement rocks of the greenland craton, in addition to some reworking of clastic sediments (surlyk 2003). the devonian–carboniferous sediments exposed north of jameson land during the early jurassic were onlapped during the middle jurassic and thus became less important sediment sources, but this situation may have changed during the late jurassic due to increased rifting and consequent exhumation and erosion of uplifted fault-block crests. samples and methods provenance analysis was performed on sandstones from six intervals in the blokelv-1 core (fig. 2). five of the samples are from the katedralen member (sample id 2–6) and one is from the sjællandselv member (sample id 1) of the hareelv formation. the sampled sandstones are structureless, grey, fineto medium-grained, moderately to well-sorted and the grains are angular to sub-rounded (bjerager et al. 2018a, this volume). some of the sampled sandstones comprise high-density turbidites (sample id 1, 4) whereas others are injectites (sample id 2, 3, 5, 6). zircon u–pb geochronology samples were crushed in order to liberate the individual grains and then sieved to retrieve the fraction <500 µm. heavy-mineral concentrates were produced on a watershaking wilfley table. zircon grains were hand-picked n deep-water marine sandstone bodies drowned (relict) sandy shelf (olympen fm, zeus mb)shallow marine sandstone hareelv fm, katedralen mb raukelv fm hareelv fm, sjællandselv mb offshore marine mudstone inferred coastline transition zone marine sandstone/mudstone heterolith 50 kma b mol prov 3 jameson land traill ø milne land fig. 3. palaeogeographic reconstructions of jameson land and surrounding areas based on surlyk (2003) and bjerager et al. (2018b, this volume). the facies distribution is tentative in most areas since the distribution in the present-day offshore area is inferred and the succession has been removed by erosion in some of the present-day onshore areas. 119 in a random fashion to ensure that all grain sizes, shapes and colours were represented. the grains were cast into epoxy and polished to expose a central cross-section of each grain. the mount was cleaned in an ultrasonic bath with propanol, and then loaded into the sample cell of the laser ablation system for age dating. approximately 160 grains were analysed per sample ensuring that even small age populations are detected (vermeesch 2004). all data were acquired with a single spot analysis on individual zircon grains using a beam diameter of 30 µm and a crater depth of c. 15–20 µm. the amount of ablated material approximates 200–300 ng for the ablation time of 30 sec. the ablated material was analysed on an element2 (thermo finnigan, bremen) single-collector, double focusing, magnetic sector-field, inductively coupled plasma mass spectrometer (icp-ms) with a fastfield regulator for increased scanning speed. the total acquisition time for each analysis was 60 sec., with the first 30 sec. used to measure the gas blank. the instrument was tuned to give large, stable signals for the 206pb and 238u peaks, low background count rates (typically around 150 counts per second for 207pb) and low oxide production rates (238u16o/238u generally below 2.5%). 202hg, 204(pb+hg), 206pb, 207pb, 208pb, 232th and 238u intensities were determined through peak jumping using electrostatic scanning in low-resolution mode and with the magnet resting at 202hg. mass 202hg was measured to monitor the 204hg interference on 204pb where the 202hg/204hg ≡ 4.36, which can be used to correct significant common pb contributions using the model of stacey & kramers (1975). the elemental fractionation induced by the laser ablation and the instrumental mass bias on measured isotopic ratios were corrected through standard-sample bracketing using the gj-1 zircon ( jackson et al. 2004). long-term external reproducibility was monitored by repeated analyses of the plešovice zircon standard (slama et al. 2008). the analytical data are reported in an online supplementary data file. the reported ages are based on 207pb/206pb derived ages for the >0.8 ga (billion years) analyses, and 206pb/238u ages for the <0.8 ga analyses, as the latter is more precise for the younger age range. the propagation of the analytical errors follows the principles of sambridge & lambert (1997). age measurements lacking a stable 207pb/206pb plateau or with a u/pb or pb/ pb error >10% were discarded. a correction for common pb was applied on up to 7% of the concordant analyses from each sample. the data are plotted using kernel density estimation (vermeesch 2012) employing analyses that are <10% discordant. heavy-mineral analysis the grain-size interval 45–710 µm was collected by sieving of crushed samples. heavy-liquid separation utilised bromoform with a specific density of 2.82 g/cm3, and the total heavy-mineral weight percentage (wt%) of the chosen grain-size interval was measured. the heavy-mineral concentrates were embedded in epoxy and polished. heavy-mineral chemistry was determined by computercontrolled scanning electron microscopy (ccsem) at geus on a philips xl40 sem equipped with two energy dispersive x-ray (edx) detectors: a thermo nanotrace and a pioneer voyager detector (keulen et al. 2008, 2012). approximately 1200 grains were analysed in each sample, of which between 22 and 327 grains were used in the mineral statistics. this is because the heavymineral samples contain mica minerals and authigenic heavy minerals, which were excluded from the heavymineral suite results due to their dependency on depositional environment and diagenesis. as there is a large compositional overlap between the minerals amphibole, pyroxene and tourmaline in standard-less edx analyses, these have been combined into one group in this study, labelled mafic minerals. the group of other heavy minerals includes ilmenite, corundum, monazite and xenotime, each of which is present in amounts <5%. results the zircon u–pb age data can be found in the online supplementary material. the age data are displayed using kernel-density estimation (fig. 4) and cumulative age distribution (fig. 5). the zircon ages cover a broad archaean to palaeozoic age span, but with the dominant zircon u–pb ages occurring between 2.0 and 1.0 ga (fig. 4). eight peak ages are present when combining all data: 2.71, 2.49, 1.95, 1.65, 1.49, 1.37, 1.10 and 0.43 ga. the relative proportion between the ages in each sample is illustrated by dividing the data into six age intervals (table 1). the grains older than 2.2 ga comprise 12– 23% of each sample, whereas grains with ages of 2.2–1.8 ga constitute 6–23%. pronounced peak ages are present within each of the 1.8–1.6, 1.6–1.3 and 1.3–0.8 ga age intervals, which cover 17–25, 14–22 and 18–34% of the measured ages, respectively (table 1). the youngest age interval of 0.6–0.3 ga is present in small proportions of 0–3% in the samples. the shallowest sample (sample id 1) is the only sample that does not have an age peak at 1.66–1.64 ga, which is pronounced in the other samples (fig. 4). the 120120 cumulative age distributions show that no significant difference exists between the six measured age distributions since the distance between the curves is small (fig. 5). kolmogorov-smirnov (k-s) tests (guynn & gehrels 2010) were performed on the zircon age data to determine if the samples were derived from different sources by comparing the distance between the cumulative agedistribution curves (fig. 5). two samples are regarded as significantly dissimilar if they have a p-value <0.05 and they will then probably have different provenance. the results show that all p-values are >0.05 (table 2). it is therefore unlikely that the samples came from different parent populations at a 95% confidence level. the samples id 2–4 resemble each other well according to the high p-values of 0.99. the sample id 5 shows good resemblance to the samples id 1 and id 6 (p-values of 0.97 and 0.91, respectively). pr ob ab ilit y 0 30 20 10 0 6 4 2 0 6 4 2 pr ob ab ilit y pr ob ab ilit y pr ob ab ilit y pr ob ab ilit y pr ob ab ilit y pr ob ab ilit y 0 6 4 2 0 6 4 2 8 0 6 4 2 0 6 4 2 age (ga)0.0 1.0 2.0 3.0 4.00.5 1.5 2.5 3.5 neoproterozoic mesoproterozoic palaeoproterozoic archaeanphanerozoic id 1 id 2 id 3 id 4 id 5 id 6 7.41 m 43.16 m 69.78 m 139.27 m 181.53 m 221.18 m n/n = 115/158 n/n = 113/156 n/n = 100/155 n/n = 118/158 n/n = 100/156 n/n = 114/162 id 1–6 all data n/n = 660/945 1.65 0.43 2.49 2.71 1.95 1.10 1.37 1.49 1.14 1.36 1.65 2.71 2.70 0.45 1.65 1.76 1.47 1.13 1.92 0.97 2.46 1.37 2.83 0.43 1.10 1.43 1.64 1.87 2.67 1.97 2.52 2.83 0.42 1.09 1.49 1.66 1.96 2.79 1.32 2.42 1.11 1.43 1.64 1.96 2.770.44 2.54 1.35 2.96 1.10 1.50 1.73 2.740.43 2.46 1.93 2.27 2.99 n um be r o f g ra in s mol prov 4 fig. 4. zircon u–pb ages of the hareelv formation. the ages are plotted with a black line and annotated peak ages using kernel density estimation (vermeesch 2012) and histograms with a 25 myr bin interval. only zircon ages with <10% discordance are plotted, which is signified by “n/n” (the number of concordant analyses out of the total number of analyses). the lowermost panel shows the distribution of all zircon ages from the six samples, and this dotted line is shown on all samples for comparison. 0.3 0.8 1.3 1.6 1.8 2.2 0.6 1.3 1.6 1.8 2.2 3.8 7.41 m id 1 2 27 20 17 12 23 43.16 m id 2 3 27 22 19 12 17 69.78 m id 3 0 34 17 24 6 19 139.27 m id 4 3 31 18 20 16 12 181.53 m id 5 id 6 3 24 14 18 23 18 18221.18 m 2 18 17 25 20 age interval (ga) min. age max. age number of zircon grains (%) table 1. distribution in six age intervals of the analysed zircon grains with concordant ages mol prov table 1 121 the heavy-mineral assemblage consists, in decreasing order of abundance, of rutile, mafic minerals, zircon, leucoxene, garnet, apatite, epidote, olivine and small amounts of other minerals (fig. 6). the garnets are classified as almandine and generally contain more mg than ca. when including mica minerals in the heavy-mineral assemblage, they constitute 60–91% of the mineral grains. muscovite comprises 72–98% of the mica minerals. the heavy-mineral content is largest in the uppermost sample (table 3). discussion the age distribution of all analysed zircon grains from the hareelv formation is shown in fig. 7a, and selected published zircon age data from older rocks in the area are shown in figs 7b–h for comparison. basement and metasediment signal the dominant zircon age populations present in the crystalline basement complexes of east greenland (fig. 1) have age ranges of about 2.8–2.5 and 2.0–1.8 ga (fig. 7h; thrane 2002). these archaean and palaeoproterozoic rocks are primarily orthogneisses of granitic to tonalitic composition, and their heavy-mineral assemblages are therefore dominated by mafic minerals and mica minerals. the metasedimentary rocks of the krummedal supracrustal sequence (fig. 1) mainly have zircon ages of 1.8–1.0 ga (fig. 7g; watt et al. 2000; leslie & nutman 2003). they consist primarily of pelitic and psammitic metasediments that underwent intermediate to highgrade metamorphism, resulting in a heavy-mineral assemblage dominated by mica minerals and garnet with localised sillimanite (watt & thrane 2001; leslie & nutman 2003). the zircon ages of the archaean and palaeoproterozoic basement complexes in east greenland correspond largely to two of the oldest age populations recorded in the hareelv formation with peak ages of 2.71 and 1.95 ga (fig. 7). however, the latest archaean to earliest palaeoproterozoic ages in the hareelv formation with a peak age of 2.49 ga are not comparable to ages reported by thrane (2002) from the basement complexes west of kong oscar fjord (fig. 1); such ages may therefore be present in other areas of the complexes. the ages of the krummedal supracrustal sequence concur with the late palaeoproterozoic and the mesoproterozoic age populations in the hareelv formation with peak ages of 1.65, 1.49 and 1.10 ga (fig. 7). the mesoproterozoic age population with peak age of 1.37 ga in the hareelv formation is only pronounced in some of the samples (fig. 4), and the population is not evident in the krummedal supracrustal sequence (fig. 7). the late palaeoproterozoic and late mesoproterozoic age populations are dominant both in the hareelv formation and in the krummedal supracrustal sequence. however, id 1 id 2 id 3 id 4 id 5 id 6 7.41 m id 1 0.447 0.382 0.477 0.966 0.476 43.16 m id 2 0.447 0.993 0.992 0.477 0.093 69.78 m id 3 0.382 0.993 0.880 0.194 0.080 139.27 m id 4 0.477 0.992 0.880 0.301 0.122 181.53 m id 5 0.966 0.477 0.194 0.301 0.909 221.18 m id 6 0.476 0.093 0.080 0.122 0.909 table 2. p-values of the k-s test of the zircon age data mol prov table 2 < 45 µm 45–710 µm > 710 µm 45–710 µm total 7.41 m id 1 5.3 23.3 71.4 4.74 1.10 43.16 m id 2 18.5 81.0 0.5 0.94 0.76 69.78 m id 3 10.8 89.2 0.0 0.87 0.77 139.27 m id 4 6.9 table 3. grain size and heavy-mineral content of the analysed sandstones 50.6 42.4 0.47 0.24 181.53 m id 5 5.4 37.5 57.1 1.21 0.45 221.18 m id 6 6.2 30.9 62.9 1.50 0.46 grain size of sandstones (wt%) heavy-mineral content (wt%) mol prov table 3 0.0 1.0 2.0 3.0 4.0 age (ga) 0.5 1.5 2.5 3.5 100 0 80 60 40 20 c um ul at ive % id 1 id 2 id 3 id 4 id 5 id 6 7.41 m 43.16 m 69.78 m 139.27 m 181.53 m 221.18 m mol prov 5 fig. 5. cumulative age distributions of zircons from the hareelv formation. 122122 the metamorphic rocks of the east greenland caledonides comprise a complex series of lithologies and the krummedal supracrustal sequence is not representative for them all. the ages of the caledonian granites that formed between 466 and 423 ma in east greenland (leslie & nutman 2003; kalsbeek et al. 2008a; rehnström 2010) are in agreement with the minor age population with a peak age of 0.43 ga present in five of the samples from the hareelv formation (fig. 4). the high contents of fe and mg in the garnets in the hareelv formation suggest an origin from a felsic al-rich rock type that has been subjected to intermediate to high metamorphic conditions. the presence of zircon, rutile, leucoxene and apatite (fig. 6) is also consistent with derivation from felsic source rocks such as granites and granitic gneisses. the proportion of mafic minerals relative to the remaining heavy minerals is smaller in the hareelv formation than in the probable crystalline and metasedimentary source rocks in the east greenland caledonides. this may be indicative of several cycles of sedimentation prior to deposition in the jameson land basin since mafic minerals break down easily during weathering and transport (morton & hallsworth 1999). furthermore, amphibole and sillimanite largely dissolve during the first 3 km of burial (andò et al. 2012) and the studied hareelv formation was buried to about 2.8 km prior to cenozoic uplift (green & japsen 2018, this volume). hence, the proportions of the mechanically and chemically stable heavy minerals, such as zircon and rutile, are high compared to the other heavy minerals in the hareelv formation (fig. 6). the negligible amount of ilmenite in the heavy-mineral suite of the sandstones probably indicates that leucoxene formed at the expense of ilmenite (weibel 2003). the interpretation of the non-mica heavy-mineral assemblage reported in fig. 6 is based on rather few grains because of the high content of mica minerals and authigenic heavy minerals (ankerite and pyrite) in the samples, so the variations in the relative proportions between the heavy minerals are not significant. it should be noted, however, that a similar suite of heavy minerals is present in all the samples (fig. 6). the provenance signal of the two finest-grained samples (id 2, 3; table 3) shows no significant deviation from the remaining samples (tables 1, 2, figs 4–6) so it is not influenced by grain size. the number of grains >710 µm in size is overestimated (table 3) since it was not possible to disintegrate all individual grains during crushing of the samples. reworking of older sediments zircon age data from the neoproterozoic eleonore bay supergroup, including the nathorst land group and the lyell land group, have peak ages of 1.75 and 1.05 ga, respectively (fig. 7; dhuime et al. 2007; slama et al. 2011). these ages are not prominent in the hareelv formation, so only a minor amount of sediment can have been reworked from this part of the eleonore bay supergroup into the jameson land basin during the late jurassic. however, it should be noted that the seven samples from the eleonore bay supergroup are not considered representative of the entire 14 km thick succession. devonian and carboniferous sediments in the kong oscar fjord area (fig. 1) have a wide range of age populations (slama et al. 2011) that resemble those in the hareelv formation (fig. 7). the relative dominance of the various age peaks, however, is not similar for the different sediments (this study and slama et al. 2011): the devonian sediments contain fewer archaean ages and more early neoproterozoic and palaeozoic ages than the hareelv formation, whereas the carboniferous sediments contain fewer middle palaeoproterozoic ages and more palaeozoic ages than the studied sandstones. however, some degree of modification of the provenance signal id 1 7.41 m n = 22 id 2 43.16 m n = 327 id 3 69.78 m n = 148 id 4 139.27 m n = 114 id 5 181.53 m n = 40 id 6 221.18 m n = 60 mafic mineral other olivineleucoxene rutile zircon epidote apatite garnet mol prov 6fig. 6. heavy-mineral contents of the hareelv formation excluding mica minerals and authigenic heavy minerals. the number of employed analyses (out of the total number of analyses of c. 1200 per sample) is indicated by ‘n’. 123 is likely to have occurred during erosion and transport. furthermore, the paucity of the database that provides the comparison should be taken into account; the 8 km thick devonian succession is covered by only three samples and the 2 km thick carboniferous succession is represented by two samples. thus, the hareelv formation may have been sourced by reworking of devonian and/or carboniferous sediments into the jameson land basin. ultimately, all the sediments originate from the same areas in the caledonides, which must comprise a combination of the archaean–palaeoproterozoic crystalline complexes (with their late archaean to middle palaeoproterozoic zircon age spectrum), the meso-neoproterozoic metasediments (with their late palaeoproterozoic to late mesoproterozoic zircon age spectrum) and the intruded caledonian granites (fig. 7). these units are closely interrelated in the present-day exposed part of the east greenland caledonides (fig. 1), where a thick section of the krummedal supracrustal sequence must have been eroded to ex0 4 2 0 40 20 60 pr ob ab ilit y pr ob ab ilit y pr ob ab ilit y pr ob ab ilit y pr ob ab ilit yn um be r o f g ra in s 0 8 4 12 0 8 4 12 0 30 20 10 0 6 4 2 0 4 2 pr ob ab ilit y pr ob ab ilit y a b c d e f g h age (ga)0.0 1.0 2.0 3.0 4.00.5 1.5 2.5 3.5 neoproterozoic mesoproterozoic palaeoproterozoic archaeanphanerozoic 1.11 1.63 2.73 0.44 1.11 1.65 2.68 0.43 1.87 1.48 1.38 0.97 1.05 1.50 2.67 1.95 1.84 0.44 1.08 1.62 1.07 1.751.64 2.73 1.11 1.51 1.67 0.42 archaean–palaeoproterozoic rocks crystalline basement complexes meso–neoproterozoic metasediments krummedal supracrustal sequence neoproterozoic sediments eleonore bay supergroup nathorst land group neoproterozoic sediments eleonore bay supergroup lyell land group middle devonian sediments kap kolthoff group upper carboniferous sediments traill ø group lower jurassic sediments neill klinter group upper jurassic sediments hall bredning group hareelv formation 6 samples, n = 660 id 1–6, this study 5 samples, n = 510 slama et al. 2011 2 samples, n = 192 slama et al. 2011 3 samples, n = 314 slama et al. 2011 3 samples, n = 339 slama et al. 2011 4 samples, n = 59 dhuime et al. 2007 3 samples, n = 62 watt et al. 2000, leslie & nutman 2003 11 samples, n = 53 thrane 2002 pr ob ab ilit y 0 30 20 10 1.65 0.43 2.49 2.71 1.95 1.10 1.37 1.49 mol prov 7 fig. 7. zircon u–pb ages of the analysed upper jurassic sandstones (a) compared to older sediments and basement complexes (b–h; watt et al. 2000; thrane 2002; leslie & nutman 2003; dhuime et al. 2007; slama et al. 2011). ages with <10% discordance are employed and they are plotted using kernel density estimation (vermeesch 2012). the locations of the samples are shown in fig. 1. 124124 pose the underlying crystalline basement (higgins et al. 2004). the liverpool land high was uplifted during the early jurassic and supplied sediment to the jameson land basin (surlyk 2003). the dominant age populations with peak ages of 1.62 and 0.44 ga found in lower jurassic sediments in the eastern jameson land basin (fig. 7; slama et al. 2011) correspond to ages present in liverpool land in the eclogite terrane and in caledonian granitoid rocks, respectively (augland et al. 2011, 2012). this distinctive age signature is not present in the hareelv formation, so the latter cannot have received material from liverpool land or its derived sediments. this is in accordance with the gradual jurassic onlap onto liverpool land (surlyk 2003) indicating that it was no longer a source area in late jurassic times. sediment transport it is plausible that the sediments in the hareelv formation and the associated shelf-edge systems were derived partly or fully from reworking of palaeozoic sediments, as described above. in that case, a contribution may have come from immediately north-west of the basin, where carboniferous sediments are exposed today (fig. 1), and/ or from farther south where these strata were probably outcropping in jurassic times. the paleozoic sediments were onlapped during the middle jurassic, but increased faulting associated with the rifting that began in the middle jurassic is known to have led to local erosion of fault crests during the late jurassic (surlyk 2003). such faulting may also have occurred inland (west) of the jameson land basin, thereby resulting in reworking of palaeozoic sediments. small differences exist between the zircon age distributions of the hareelv formation (table 1; fig. 5), but none of the analysed samples was found to be statistically dissimilar according to the performed k-s tests (table 2). this suggests a consistent and persistent transport path for the sands deposited in this part of the jameson land basin in the late jurassic. the results of the provenance analysis of the hareelv formation may also be used to characterise the shelfedge sediments that collapsed and subsequently were redeposited by gravity flows (fig. 3). the feeding shelfedge sands are represented by the proximal facies of the olympen formation (or younger equivalents) during deposition of the katedralen member and subsequently by the raukelv formation during deposition of the sjællandselv member (surlyk 2003; surlyk et al. 2007; bjerager et al. 2018b, this volume). the olympen formation (zeus member) and the raukelv formation (fig. 2) are both interpreted to have been deposited during relative lowstands in the region, but the source terranes that supplied sediment to the northern shelf edge of the jameson land basin may not have been entirely the same during these two time intervals since there are some distinctive differences between the samples from the katedralen and sjællandselv members. the sandstone from the sjællandselv member (id 1) has a higher heavy-mineral content (table 3) and a higher proportion of archaean zircon grains (table 1) than the sandstones from the katedralen member (id 2–6). the sjællandselv member sample has a peak age at 1.73 ga, but lacks the pronounced peak age at 1.66–1.64 ga in the remaining samples (fig. 4), and sandstones from the sjællandselv member are coarsergrained than the katedralen member sandstones (surlyk 2003; bjerager et al. 2018a, this volume). furthermore, the sjællandselv member sample contains 4% rock fragments compared to only 0–1% in the samples from the katedralen member (olivarius et al. 2018, this volume). this could indicate rejuvenation of the topography during the volgian such that the raukelv formation received some of its sediment from more proximal sources in comparison to the previous shelf-edge sands. this is in accordance with the late jurassic rifting that caused block faulting and thereby created local sediment sources through erosion of tilted footwall crests (surlyk 2003). conclusions the large range of zircon u–pb ages in the upper jurassic hareelv formation resembles the zircon age spectrum in the caledonides of east and north-east greenland. the heavy-mineral suite and garnet composition in the sandstones indicate derivation from felsic rocks. archaean–palaeoproterozoic crystalline complexes, meso– neoproterozoic metamorphic rocks and caledonian migmatites and granites are considered to represent the original source of the material. some or all of the material may, however, have undergone several sediment cycles before deposition in the jameson land basin in the late jurassic. sediment supply to the hareelv formation from reworking of devonian and carboniferous sediments is plausible since the preserved parts of these successions contain corresponding age populations. however, it should be emphasised that the database for comparison is limited. 125 the provenance signature is a reflection of the area exposed to lowstand erosion during the accumulation of the southwards prograding shelf-edge deltas that subsequently fed the gravity flows of the hareelv formation. the provenance signature is rather uniform in the oxfordian–kimmeridgian katedralen member, but there are a number of indications of a shorter sediment transport distance for the sands of the early volgian sjællandselv member, potentially reflecting rejuvenation of the topography during rifting. supplementary data to this article can be found online at: www.geus.dk/bulletin42 acknowledgements john boserup, anders pilgaard, fiorella f. aguilera, mojagan alaei, jørgen kystol and ingerlise nørgaard are thanked for sampling and analysis. dirk frei, niels henriksen, kristine thrane and peter appel provided valuable advice and jette halskov kindly did the artwork. the referees, jiri slama and martin sønderholm, are thanked for insightful comments that significantly improved the manuscript. references andò, s., garzanti, e., padoan, m. & limonta, m. 2012: corrosion of heavy minerals during weathering and diagenesis: a catalog for optical analysis. sedimentary geology 280, 165–178. augland, l.e., andresen, a. & corfu, f. 2011: terrane 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the geological society of denmark 50, 171–183. _________________________________________________________________________________ manuscript received 18 december 2015; revision accepted 29 august 2017 geological survey of denmark and greenland bulletin 26, 2012, 81-84 81© 2012 geus. geological survey of denmark and greenland bulletin 26, 81–84. open access: www.geus.dk/publications/bull the cenozoic song hong and beibuwan basins, vietnam michael b.w. fyhn, henrik i. petersen, lars henrik nielsen, tran c. giang, le h. nga, nguyen t.m. hong, nguyen d. nguyen and ioannis abatzis the vietnamese offshore margin holds a substantially underexplored petroleum potential. the key to unravelling this potential lies in understanding the tectono-stratigraphic framework of the region including the cenozoic mechanisms governing syn-rift and source rock deposition. this is essential for prediction of, for instance the presence and nature of source rocks in south-east asia and possible reservoir intervals in the syn-rift packages. the vietnamese part of the song hong and beibuwan basins (fig. 1) differs from other basins along the western margin of the south china sea in that the palaeogene syn-rift succession is sporadically exposed due to uplift and inversion. these exposures provide a unique glimpse into the cenozoic syn-rift succession of the basin. considerable oil and condensate discoveries have recently been made in the vietnamese part of the song hong basin. this emphasises the need for improved geological models of the basin in order to unravel the local petroleum systems. since 1995, the geological survey of denmark and greenland and the vietnam petroleum institute have conducted exploration-related, foreign development aid projects in vietnam funded by the danish ministry of foreign affairs. the ongoing danish–vietnamese research cooperation operates through an enhanced research capacity (enreca) project. the enreca group consists of both danish and vietnamese researchers who, through the project, have become the established experts in the region. the ongoing enreca work focuses on the geological development of the greater song hong basin (including the vietnamese part of the beibuwan basin). as part of the project activities, the enreca group has recently drilled a fully cored well into the syn-rift succession of the basin and is carrying out a comprehensive regional study combining seismic, well and outcrop information. evolution of the the song hong basin and the greater vietnamese margin south-east asia is tectonically one of the most complex regions in the world, shaped by continental collisions and the creation and subduction of oceanic basins. a large portion of the deformation associated with the indian–eurasian plate collision was likely accommodated in the neighbouring part of south-east asia during the cenozoic (e.g. leloup et al. 2001). indochina deformed in response to the collision and may have been squeezed hundreds of kilometres south-eastcuu long b asi n malay basin m ae ping shear zone pearl r iver mouth basin w. natuna basin n am c on s on b as in dangerous grounds malay basin china borneo basins with palaeogene lacustrine and coaly syn-rift source rocks oceanic crust with numbered main cenozoic strike-slip direction 6 m al ay si a cambodia beibuwan basin bach long vi qiongdongnan basin phu khanh basin song h ong basin laos thailand 11 11 11 11 10 10 9 9 8 8 5e 5e 5e 6a 6a 5e 5e6a 6b 6b 6b 6b 6a 6a 6 6 6a 6b vietnam 110°e 20°n 5°n 250 km red river fault zone magnetic anomalies nha trang quang ngai south china sea fig. 1. map of indochina and the western part of the south china sea showing the outline of the main cenozoic structures and basins with lacustrine and coaly source rocks. the large red rectangle shows the study area, and the small rectangle shows the location of fig. 3. 8282 wards from the himalayan collision front along enormous strike-slip fault zones (fig. 1). the song hong basin is the largest vietnamese sedimentary basin and covers close to one third of the western margin of the south china sea. its formation has been attributed to large-scale, left-lateral movements along the vietnamese margin (e.g. leloup et al. 2001), whereas a pull-apart model for the basin remains controversial (e.g. hall & morley 2004). understanding the basic mechanism behind the formation of the song hong basin is crucial for understanding the overall establishment of the western half of the south china sea. the song hong basin is likely to hold important information on the opening mechanism of the south china sea. whether the greater vietnamese margin constitutes a transform margin or a more classic oblique rift margin is fundamental to the structural concepts applied when interpreting geophysical data across basins along the margin. so far it has not been possible to evaluate this satisfactorily due to the scarcity of seismic data tied to a limited number of wells offshore vietnam. however, recent studies of relatively densely spaced seismic profiles tied to wells offshore central and southern vietnam have led to an increased understanding of the tectonic evolution of the margin (fyhn et al. 2009). it is clear from these studies that the song hong basin is crucial for reconstructing the overall tectonic development of the western south china sea. the current enreca study is based on a comprehensive geoscientific database of the vietnamese portion of the basin and aims to unravel the tectonic and depositional evolution of the basin in a regional context. this will allow a detailed stratigraphic understanding of the basin fill and a comprehensive structural analysis of the basin and the adjacent areas. the song hong basin forms a depocentre in the northwestern part of the south china sea and is filled with up to 15–20 km of sediments. sediment dispersal to the northwestern part of the south china sea has largely occurred as spill from the song hong basin. the distribution and style of sediments in the surrounding basins are therefore closely linked to the depositional pattern in the song hong basin. deposition evolved rapidly in the basin and varied in time and space from lacustrine and fluvial to marine sedimentation. sedimentary environments ranged from carbonate platform, siliciclastic and lava deltas to shelf, slope and basin floor and were subject to relative sea-level changes. combined with high sedimentation rates this makes the basin well suited for studies of cenozoic depositional trends along the western south china sea. the depositional development in the basin is investigated through seismic stratigraphic and facies analyses, and well data are closely integrated in the analysis in order to extract age and lithologic information. palaeogeographic reconstructions illustrate the gross-sedimentary evolution of the basin and form a primary result of this part of the study (fig.  2). information on the regional geological evolution can be obtained from the depositional style of the basin in addition to information on the establishment of the south i i i i i i i iiii i i i i i i i i i i i i i i i iiii i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i ii i i i i i i i ii i i i i i i ii i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i ii i i i i i i i i i i i i i ii i i i i i i i i i i i i i i i i i i i i i i i i i i i i ii i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i iiiii i i i i i i i i i i i i i ii i i i i i i i i i i i i i i i i i i i i i ■■ quang ngai 100 km intra-miocene nha trang tuy hoa qui nhon carbonate platform (predominant) carbonate detritus apron carbonate ramp alluvial – neritic (siliciclastics) bathyal (siliciclastics) slope (siliciclastics) volcano/basalt delta geology concealed underneath volcanics subaerially exposed non-deposition present land with major cities steep platform margin active fault i i i i i 109°e 13°n 13°n fig. 2. intra-miocene depositional map for the southern part of the song hong and the phu khanh basins. carbonate deposition took place on structural highs, but was affected by siliciclastic input sourced from onshore uplift and local volcanism. 83 china sea, regional uplift history, sediment transport and cenozoic climate. neogene basalt volcanism affected large parts of the vietnamese margin from the nam con son basin in the south to the song hong basin in the north, and acted coeval with widespread carbonate platform growth along the margin (fig. 2). magmatism was associated with regional onshore uplift and denudation, which seems to have had a significant derivative effect on contemporary deposition (fyhn et al. 2009). increased siliciclastic input to the offshore basins and elevated inorganic nutrient concentrations in platform areas associated with onshore uplift and denudation led to a lowering of long-term carbonate production and deposition. as a result, carbonate platforms drowned and became buried underneath kilometre-thick, prograding siliciclastic sediment wedges during the late neogene. in the southern song hong basin, volcanics and carbonate platforms are nicely revealed by seismic data and are drilled and dated in wells (fig. 2). syn-rift and source-rock deposition cretaceous and cenozoic organic-rich lacustrine mudstones and humic coals, which were deposited during periods of rifting, form world-class source-rock intervals in south-east asia (fig.1). these source rocks frequently occur adjacent to coarse-grained, reservoir-quality sedimentary bodies derived from nearby elevated hinterland areas. moreover, source-rock intervals are often located adjacent to fractured basement or pre-rift successions with effective reservoir characteristics. understanding these rift systems, including the occurrence and genesis of lake systems and the sedimentary successions capable of generating and storing hydrocarbons, is critical when evaluating local petroleum systems over larger parts of south-east asia. it has long been recognised that stratigraphic wells through syn-rift successions are useful to unravel the deeper depositional systems including source-rock units (e.g. sladen 1997). this is particularly true offshore northern vietnam, where syn-rift lacustrine and coaly-sourced petroleum mainly occur in fractured and karstified pre-rift carbonates. even so, little is known about the source-rock system, and deep wells drilled into the syn-rift succession have not provided information on the source rocks of the area. to address this problem, the enreca group drilled a 500 m fully cored well into the inverted syn-rift succession on the island of bach long vi (figs 1, 3) where highly oil-prone source-rock intervals are widespread. the core demonstrates a thick lacustrine succession within the area. at the time of writing, cores are transported to a core depository to be analysed in detail. bach long vi has low relief and dense vegetation, and hence well-exposed outcrops on the island are limited to locations along the shore, and a satisfactory study of syn-rift sedimentation and source-rock quality is not possible based on outcrops alone. however, combined with the cored stratigraphic well, the geology of bach long vi provides a unique opportunity to study syn-rift deposition and source-rock composition. the core provides a regional type section of i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i iiiiiiii i i i i i i i i i i i i i i iiiii i i ii i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i iii ii iii b do son yen tu chi linh tien lang 20 km palaeogene rift system n bach long vi fig. 3a nw se 0 500 1000 1500 post-rift inverted syn-rift 5 km bach long vi well projected c. 5 km along structural strike a t w o -w ay t ra ve l ti m e (s ec ) major fault seismic line i i i neogene eocene–oligocene syn-rift basement and palaeozoic–mesozoic pre-rift basement/syn-rift prospects fig. 3. a: geotransect across the inverted palaeogene graben, which is exposed on the island of bach long vi a few kilometres from the profile. the approximate well trajectory is projected onto the section. the transect is based on the seismic line 89-1-103 (partly shown on fig. 3b, red, dotted line). b: map showing the outline of the primary paleogene syn-rift depocentre and syn-rift structural highs in the north-eastern song hong basin. oil and condensate discoveries in the area are sourced from lacustrine and coaly syn-rift source rocks. bach long vi forms a window to a graben directly linked with the syn-rift kitchen areas of the recent discoveries made in fractured and karstified palaeozoic carbonates in the area. 8484 the palaeogene syn-rift succession. sampled source-rock intervals are studied through general source rock screening and biomarker analysis combined with more advanced analysis of selected source-rock intervals. situated at the intersection of the beibuwan and the song hong basin, bach long vi occurs in the centre of a newly discovered petroleum province offshore vietnam (fig. 3). bach long vi sits at the top of a prominent structure associated with inversion of a palaeogene graben. the island therefore provides a window into the nature of the palaeogene syn-rift succession of the area. highly oil-prone lacustrine mudstones with an organic content ranging from c. 2 to 7 wt.% total organic carbon were sampled from scattered outcrops on bach long vi and from sub-sea outcrops around the island (fig. 4). enreca analyses of these excellent source rocks confirm that petroleum-prone mudstones occur at various stratigraphic levels that crop out over the entire area (e.g. petersen et al. 2004 and new unpublished data). a dense net of seismic data cover the area offshore bach long vi and reveals a n–s to ne–sw-trending, inverted eocene–oligocene graben. the graben fill sub-crops at and closely beneath the seafloor and forms the core of bach long vi (fig. 3a). the graben is part of a larger rift system that actively subsided during palaeogene time (fig. 3b). the deposits, including the sampled source rocks on and around bach long vi, therefore form part of a regional syn-rift system, which is directly continuous with the syn-rift kitchen areas sourcing the recent oil and condensate discoveries in the region (fig. 3). the preliminary results suggest that rich source-rock intervals could be present throughout much of the syn-rift succession in the basin, highly encouraging for further exploration. acknowledgements the enreca project is funded by the danish ministry of foreign affairs via danida. petrovietnam co-funded the enreca 3 well and is thanked for providing data and for permitting publication of this paper. references fyhn, m.b.w., boldreel, l.o. & nielsen, l.h. 2009: geological development of the central and south vietnamese margin: implications for the establishment of the south china sea, indochinese escape tectonics and cenozoic volcanism. tectonophysics 478, 184–214. hall, r. & morley, c.k. 2004: sundaland basins. in: clift, p. et al. (eds): continent–ocean interactions within the east asian marginal seas. geophysical monograph series 149, 55–85. washington d.c.: american geophysical union.leloup, p.h., arnaud, n., lacassin, r., kienast, j.r., harrison, t.m., trong, t.t.p., replumaz, a. & tapponnier, p. 2001: new constraints on the structure, thermochronology and timing of the ailao shan-red river shear zone, se asia. journal of geophysical research 106, b4, 6683–6732. petersen, h.i., nytoft, h.p. & nielsen, l.h. 2004: characterisation of oil and potential source rocks in the northeastern song hong basin, vietnam: indications of a lacustrine–coal sourced petroleum system. organic geochemistry 35, 493–515. sladen, c. 1997: exploring the lake basins of east and southeast asia. in: fraser, a.j., matthews, s.j. & murphy, r.w. (eds): petroleum geology of southeast asia. geological society special publications (london) 126, 49–76. 20°14´n 20°13´n bach long vi 107°72´n 107°74´n500 m bach long vi 5 km fig. 4. map of the island of bach long vi. red dots show sites of sourcerock samples. the inset map shows the nearest industrial seismic lines in the area. for location see figs 1 and 3. authors’ addresses m.b.w.f., h.i.p., l.h.n. & i.a., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: mbwf@geus.dk t.c.g. & l.h.ng., vietnam petroleum institute, 173 trung kinh str., yen hoa wrd., cau giay dist., hanoi, vietnam. n.t.m.h., hanoi university of mining and geology, dong ngac, tu liem, hanoi, vietnam. n.d.n., hanoi university of science, 334 nguyen trai-thanh xuan-hanoi, vietnam. geological survey of denmark and greenland bulletin 28, 2013, 69-72 69 darkening of the greenland ice sheet due to the meltalbedo feedback observed at promice weather stations dirk van as, robert s. fausto, william t. colgan, jason e. box and the promice project team* fig. 1. map of greenland showing the locations of the 21 promice weather stations in eight regions. the blue colours show mean satellite (modis) derived albedo for the months of june, july and august for 2008–2012. the contour lines show elevations (m) of the ice-sheet surface. © 2013 geus. geological survey of denmark and greenland bulletin 28, 69–72. open access: www.geus.dk/publications/bull * andreas p. ahlstrøm, signe b. andersen, morten l. andersen, charalampos charalampidis, michele citterio, karen edelvang, trine s. jensen, signe h. larsen, horst machguth, søren nielsen, martin veicherts and anker weidick the greenland ice sheet is losing mass (barletta et al. 2012) and at least half of this loss is caused by an increase in surface melt (e.g. tedesco et al. 2013). the other part is caused by increased dynamic mass loss, as marine-terminating glaciers lose resistive stresses (nick et al. 2009) due to both retreat and meltwater lubrication at the bed (sasgen et al. 2012). in 2007, the programme for monitoring of the greenland ice sheet (promice) was initiated with the aim of gaining an insight into the causes of the ice-mass budget changes based on quantitative observations. this is primarily done by assessing how much mass is gained as snow accumulation on the surface versus how much is lost by calving and surface ablation (ahlstrøm et al. 2008). promice monitors the surface mass balance by means of automatic weather stations (awss) designed to quantify accumulation and ablation, as well as the specific energy sources contributing to ablation. these observations are vital to interpreting the physical mechanisms for ice-sheet response to climate change and for the calibration and validation of both satellite observations and climate models. in the wake of several record-breaking warm summers – increasing surface melt rate and extent (nghiem et al. 2012) – interest in greenland’s surface mass balance has increased (tedesco et al. 2013). observations of net ablation at promice stations provided in situ confirmation of extreme massloss events in 2010 (fausto et al. 2012) and 2012, primarily documented by other workers through satellite data. in this paper, we present atmospheric temperatures and surface solar reflectivity (known as albedo) of the greenland ice sheet in the promice period. albedo modulates the absorption of solar radiation, which is the primary source of melt energy. it is reported to be decreasing in greenland in recent years (box et al. 2012), causing the monitoring of albedo variability to be increasingly important. air temperatures, besides being strongly correlated to surface melt rates, affect surface albedo by controlling the rate of snow-grain metamorphism and the fraction of summer precipitation falling as rain versus snow. to elucidate the so-called melt-albedo feedback, whereby increased melt darkens the ice sheet and further enhances melt, the relationship between albedo and air temperature, observed at promice stations, is examined in this study. 1500 2000 2500 20 00 0.8 1.0 albedo 0.7 0.75 0.5 0.6 0.4 0.3 25 00 20 00 15 00 kan scoupe thu kpc tasnuk qas 3000 500 km 7070 promice measurements the original promice network consisted of 14 awss in the regions kronprins christian land (kpc), scoresbysund (sco), tasiilaq (tas), qassimiut (qas), nuuk (nuk), upernavik (upe) and thule (thu), each region monitored with a lower (l) and an upper (u) station in the ablation area (fig. 1). promice has both contributed to and received contributions from other projects in the kangerlussuaq, nuuk and tasiilaq regions, leading to the installation of seven additional stations. the promice study regions were selected to best complement the spatial distribution of existing ice-sheet weather stations, such as in the greenland climate network (steffen et al. 1996), by providing data from the under-represented ablation area (ahlstrøm et al. 2008). the promice awss measure meteorological variables including air temperature (at c. 2.7 m above the surface), pressure and humidity, wind speed, downward and upward solar (shortwave) and terrestrial (longwave) radiation. the awss also record temperature profiles in the upper 10 m of the ice, gps-derived location and diagnostic parameters such as station tilt angles. a pressure transducer and two sonic rangers measure snow and surface-height change associated with ablation and accumulation (fausto et al. 2012). most variables are measured every ten minutes, with the data stored locally awaiting collection during maintenance visits. hourly averages of the most transient variables are transmitted via satellite between days 100 and 300 of each year, while the remaining variables are transmitted at six-hour intervals. transmissions have a daily frequency in the remaining (winter) period. all data and metadata including sensor specifications are archived in the promice database and made freely available for display and download at www.promice.dk. in this study, we present monthly mean measurements of air temperature and surface albedo. to obtain the temperature averages, we first calculate daily mean air temperatures for all days in which data coverage of hourly mean values exceeds 80%. subsequently, we calculate monthly mean air temperature for all months in which data coverage of daily mean values exceeds 24 days. to calculate surface albedo, we divide instantaneous values of upward shortwave radiation by the downward component before averaging. however, instrument tilt induces significant errors in the measurement of shortwave radiation (van den broeke et al. 2004), which is a common problem in the ice-sheet ablation area due to irregular melting of the ice surface on which the aws stands. therefore we employ the tilt correction method as described by van as (2011) that uses the measured aws tilt to correct downward shortwave radiation. in contrast to the temperature averaging, a minimum of one successful hourly mean albedo is sufficient to produce a daily mean, provided that the direct solar radiation (which occurs when skies are not overcast) hits the upper dome of the radiometer at angles exceeding 30°, where measurements are more precise. the low sun angle in winter prevents calculation of albedo values. atmospheric temperature all promice sites record a distinct annual cycle in air temperature (fig. 2a). as is common for arctic climates, temporal variability is largest in winter due to a more vigorous atmospheric circulation. the amplitude in the annual air temperature cycle is largest for stations at high latitudes or high elevations since above-freezing temperatures and thus a melting ice surface capable of local thermo-regulation, are least common at these stations. the more northerly stations also show a larger annual temperature cycle due to the increasing contrast in the lengths of polar day and night with increasfig. 2. a: monthly mean air temperatures at the 18 promice sites installed on the ice sheet and before 2012. b: same but for albedo. c: albedo versus temperature. black lines: kpc stations, green: sco, orange: tas, light blue: qas, dark blue: nuk, red: kan, purple: upe, yellow: thu. year 20082007 2009 2010 2011 2012 a b c 20082007 2009 2010 2011 2012 a ir te m pe ra tu re (° c ) 10 0 –10 –20 –30 su rf ac e al be do 0.8 0.6 0.4 0.2 –25 –20 –15 –10 –5 0 5 10 air temperature (°c) su rf ac e al be do 0.8 0.6 0.4 0.2 71 ing latitude. the smallest amplitude in the annual temperature cycle is seen at qas_l, the most southerly promice site. here, free-atmospheric temperatures can exceed 20°c, leading to strong melt. in 2010, the large heat content of the atmosphere, low summer albedo and anomalously low winter accumulation combined to yield a long net-ablation season and the largest ever recorded ablation in greenland in a single melt season (9 m of ice; fausto et al. 2012). as the greater part of mass loss by melt takes place during the summer, we calculated the fouror five-year trends in combined mean air temperature for the months of june, july and august, for eight stations with a sufficiently long coverage. given the relatively short promice record length, these trends are not free from the influence of inter-annual, natural climatic variability and thus not climatological trends. the data show that at all except two sites the summers have become warmer over the promice period. most noteworthy is that the warming is most pronounced (c. 0.6°c/year) at the high latitude/elevation sites, where the influence of a melting ice surface is spatially and temporally limited. at the sites that experience the highest temperatures and strongest melt, interannual variability of free-atmospheric temperatures has had limited effect on air temperatures over the nearly permanently melting ice-sheet surface in summer. the inter-annual temperature variability is shown in an anomaly plot (fig. 3a). it is seen that 2010 was a warm year, especially in west greenland, and mostly early and late in the year, with anomalies exceeding 5°c. the widespread and large melt in 2012 (nghiem et al. 2012) was the result of high temperatures in july, as seen from positive air temperature anomalies at all promice sites. the record-warm years/summers of 2010 and 2012 (e.g. tedesco et al. 2013), following the warmest decade in greenland’s instrumental temperature record, are consistent with persistent warming observed globally, but are suggested to be a consequence of north atlantic oscillation variability affecting atmospheric heat transport (fettweis et al. 2013). atmospheric warming has been reported to be highest in south and west greenland. this is confirmed by the promice observations in west greenland, but observations at kpc_u in north-eastern greenland show a similar rate of short-term warming. these observations provide in situ indications that the atmospheric warming may be spatially pervasive. darkening ice-sheet surface the surface albedo is generally high in the cold, snow-covered interior of the ice sheet (>0.75), and lower along the ice-sheet margin where melting occurs in summer (fig. 1). in winter, the ice sheet is fully snow covered except where wind erosion dominates. depending on the location of each aws in the ablation area, snow melt starts in april or may as seen from air temperatures and decreasing albedo (fig. 2b). thereafter, albedo drops throughout the melt season until snowfall occurs in autumn, yielding a distinct annual cycle which is largest at the high-melt sites. surface melt causes this annual darkening of the ice-sheet surface as snow undergoes heatdriven metamorphosis, or completely melts to expose darker bare ice. the ice-sheet surface may also darken as impurities collect on the ice surface or supraglacial meltwater-filled features become more abundant. we find that on average surface albedo drops below fresh snow values as monthly mean temperatures exceed c. –2°c (fig. 2c). the hyperbolic shape of the scatter plot is primarily a consequence of the annual cycle in albedo (α) and can be approximated by where t is near-surface air temperature and maximum (αmax) and minimum albedo (αmin) are prescribed as 0.8 and 0.2, respectively. t0 and c, taken as 2 and 1°c respectively, are empirical constants characterising a melting point offset and the exponential scaling length of α(t). fig. 3. a: monthly mean air temperature anomalies, i.e. after subtracting the mean annual cycle for stations with data series spanning a minimum of three years. b: albedo versus temperature anomalies plotted for months with mean temperatures exceeding –2°c in order to extract melt-season values only. year –4 –2 0 2 4 air temperature anomaly (°c) a ir te m pe ra tu re a no m al y (° c ) 5 0 –5 –10 su rf ac e al be do a no m al y 0.2 0 –0.4 –0.2 r = –0.59 a b 20082007 2009 2010 2011 2012 7272 though there are exceptions, taken as a whole the promice data indicate that albedo has decreased over the past five years while temperature has increased. this is most notable at the higher-elevation sites; the lower sites are completely snow-free in every summer and thus exhibit little change. by calculating the temporal correlations between temperature and albedo for all individual stations for the months of june, july and august separately (minimum four-year time series), we isolate the inter-annual variability by eliminating the annual cycle. we observe 36 out of 39 correlation coefficients to be negative, implying a widespread association of temperature-induced melt with surface albedo on the ice sheet. mean correlation is strongest in june (–0.76 ± 0.28), followed by august (–0.54 ± 0.39) and july (–0.43 ± 0.45). the surface albedo at most awss was relatively low in 2010 and 2012, coincident with the warm summers of the past years. in order to assess the impact of atmospheric warming on ice-sheet darkening at any given site on the greenland ice sheet, we plot the albedo anomalies versus the temperature anomalies in fig. 3b for monthly temperatures exceeding –2°c, hereby isolating the melt season. the correlation of –0.59 between the plotted variables is statistically significant. a linear fit yields that one degree of warming in the near-surface air temperature will lead to an average albedo reduction of 0.043. this value is sensitive to the temperature cut-off value (here taken at –2°c) and will become more accurate with longer aws time series. since the promice stations measure close to the ice surface where temperature variability is dampened over a melting surface, a stronger correlation could be expected between albedo and local freeatmospheric temperatures, although the regions with low temperature variability are also associated with low albedo variability (fig. 3b). as mentioned, this darkening is likely due to an increase in surface melt, which cannot be linked to changes in solar radiation in recent years and thus may very well be directly and indirectly caused by changes in air temperature. although absorbed solar radiation is the primary source of melt energy, the melt-albedo feedback is initiated by the energy fluxes that respond to changes in temperature, such as downward longwave radiation and the turbulent heat fluxes. since both atmospheric warming and ice-sheet darkening increase surface melt intensity and melt area, the anticipated future warming will result in a self-reinforcing ice sheet mass-loss contribution from the melt-albedo feedback. while increased surface melt is a primary mechanism for ice loss in greenland, an increase in meltwater may also enhance mass loss due to ice dynamics, through processes such as basal lubrication and warming of the ice matrix. acknowledgements promice is funded by the danish ministry of climate, energy and building. the awss in the kangerlussuaq region are funded by the greenland analogue project, while partner awss in the nuuk region are co-funded by the greenland climate research centre. several other projects also contributed to promice. references ahlstrøm, a.p. & the promice project team 2008: a new programme for monitoring the mass loss of the greenland ice sheet. geological survey of denmark and greenland bulletin 15, 61–64. barletta, v.r., sørensen, l.s. & forsberg, r. 2012: variability of mass changes at basin scale for greenland and antarctica. the cryosphere discussions 6, 3397–3446. box, j.e., fettweis, x., stroeve, j.c., tedesco, m., hall, d.k. & steffen, k. 2012: greenland ice sheet albedo feedback: thermodynamics and atmospheric drivers. the cryosphere 6, 821–839. fausto, r.s., van as, d. & the promice project team 2012: ablation observations for 2008–2011 from the programme for monitoring of the greenland ice sheet (promice). geological survey of denmark and greenland bulletin 26, 73–76. fettweis, x., hanna, e., lang, c., belleflamme, a., erpicum, m. & gallée, h. 2013: important role of the mid-tropospheric atmospheric circulation in the recent surface melt increase over the greenland ice sheet. the cryosphere 7, 241–248. nghiem, s.v., hall, d.k., mote, t.l., tedesco, m., albert, m.r., keegan, k., shuman, c.a., digirolamo, n.e. & neuman, g. 2012: the extreme melt across the greenland ice sheet in 2012. geophysical research letters 39, l20502, http://dx.doi.org/10.1029/2012gl053611. nick, f.m., vieli, a., howat, i. & joughin, i. 2009: large-scale changes in greenland outlet glacier dynamics triggered at the terminus. nature geoscience 2, 110–114. sasgen, i., van den broeke, m.r., bamber, j.l., rignot, e., sørensen, s.l.s., wouters, b., martinec, z., velicogna, i. & simonsen, s.b. 2012: timing and origin of recent regional ice-mass loss in greenland. earth and planetary science letters 333, 293–303. steffen, k., box, j.e. & abdalati, w. 1996: greenland climate network: gc-net. in: colbeck, s.c. (ed.): glaciers, ice sheets and volcanoes: a tribute to mark f. meier. crrel special report 96-27, 98–103. tedesco, m., fettweis, x., mote, t., wahr, j., alexander, p., box, j. & wouters, b. 2013: evidence and analysis of 2012 greenland records from spaceborne observations, a regional climate model and reanalysis data. the cryosphere 7, 615–630. van as, d. 2011: warming, glacier melt and surface energy budget from weather station observations in the melville bay region of northwest greenland. journal of glaciology 57(202), 208–220. van den broeke, m., van as, d., reijmer, c. & van de wal, r. 2004: assessing and improving the quality of unattended radiation observations in antarctica. journal of atmospheric and oceanic technology 21, 1417–1431. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: dva@geus.dk untitled 90 kr03 thrusting. its relationship to the footwall flat of kr02 indicates that it was thrust along a hanging-wall flat in the order of 60 m. however, before the kr03 thrusting was complete, the kr04 thrust sheet was already emplaced on its back. the displacement of kr04 can be determined in the cross-section to be 66 m. the implications of kr04 being thrust onto kr03 are that the sedimentation of the rubjerg knude formation on top of kr03 ceased, and with the continued propagation of kr03 over the footwall ramp, the back-thrust splay also affected the kr04 thrust sheet that was being passively transported piggyback on kr03. brede rende section for more than a century, groundwater drainage has been concentrated at a spring at brede rende. from the spring, situated at the base of the cliff, a stream has over the years eroded a large funnel-shaped gully behind the cliff facing the sea. groundwater erosion successively stepping backwards is thus responsible for the wide gully and for the locality name (danish: brede = wide; rende = gully). the groundwater transmissivity is, of course, governed by the geology of the brede rende section, such that the spring wells out from the unconformity surface between the clayey lønstrup klint formation and the permeable sand of the rubjerg knude formation. the water initially drained in a southerly direction, but the present northerly drainage system is exposing the structures of the northern flank of brede rende in an isolated cliff. it is likely that this cliff will be completely removed by erosion by the sea as well as by the stream within the next few years leading to the formation of a broad gully at this location. important structural features currently exposed in the brede rende section comprise a polydiapiric complex, the brede rende diapir in the frontal part, the prominent brede rende normal fault (brnf) in the central part, and a series of duplexes stacked in the trailing end of the section. tectonic architecture the brede rende section comprises eight thrust sheets, annotated br01–br08 (plate 2). the southern and frontal boundary of the section is the footwall ramp of kr04. the northern boundary is the thrust fault which partly acts as the br08 footwall ramp and flat, and partly is the hanging-wall ramp for the sr01 thrust sheet in the sandrende section. the br01 thrust sheet is a relatively thin sheet that was displaced up along the kr04 footwall ramp, which is c. 30 m thick and dips about 35°n. above br01, the br02 thrust sheet was displaced more than 50 m along its hanging-wall ramp onto the upper footwall flat of kr04. the lower part of the thrust separating br01 and br02 has been destroyed by penetrating diapirism, and together with br03 these frontal thrust sheets in the brede rende section constitute the brede rende diapir (figs 73, 74). the thrust sheets can still be regarded as individual coherent elements, although their fig. 74. detail of the internal structure of the brede rende diapir. although the structure appears as a chaotic mixture of disrupted sand beds ‘floating’ in a disorganised fashion in the mobilised mud-matrix, some of the features could be interpreted as relicts of hanging-wall anticlines (see dashed lines) formed in a developed stage during thrusting up along steep ramps. photograph: june 1985. 91 boundaries and internal structure have been strongly distorted by diapirism. br03 is the longest thrust sheet in the brede rende section, when the trailing lower segment is included (see plates 1, 2). this is a constructional convention based on the consideration of which of the hangingwall ramps should be traced down to the décollement surface, and thus determine the annotation of the subsurface duplex sheets (plate 2, see later). the br04 thrust sheet is about 300 m long and is displaced by the brede rende normal fault (brnf) (fig. 75). north of the brnf, the br04 thrust sheet was thrust along an intermediate br03 footwall flat, and south of brnf the upper hanging-wall ramp and flat of br04 were thrust over the upper footwall flat of br03. the amount of displacement along thrust faults in this part of the brede rende section is 50 m (measured in the crosssection of plate 1) for br03 as well as br04. in br04, the rubjerg knude formation reaches its maximum thickness of about 20 m in the brede rende section, whereas the cover of rubjerg knude formation on the back of br02 and br03 is less than 5 m thick. the br05 thrust sheet is relatively short and located between br04 and br06. the displacement along its hanging-wall ramp is about 80–90 m and the initial ramp-angle was c. 12°. the thickness of the rubjerg knude formation on top of br05 is only about 5 m, which indicates that the thrusting of the br06 thrust sheet propagated early in the thrust development of the brede rende section. the piggyback thrusting of br06 on br05 on br04 is one of the best examples of a duplex structure in the rubjerg knude glaciotectonic complex. the br06 thrust sheet is a relative long and thin thrust sheet. to the north, the trailing end of br06 was thrust up over the footwall ramp of br05. from the footwall ramp hinge, an upper hanging-wall flat (br06hwf) was displaced along the upper br05 footfig. 75. the brede rende normal fault (brnf). the frontal part of the br06 thrust sheet has a normal displacement of about 20 m down through the 45° dip normal fault, which can be measured from the hanging-wall flat of br06 (br06hwf ) north of the normal fault to the br06hwf south of the normal fault. in the footwall block of the brnf, a series of minor normal faults make a stepwise displacement of the downthrown hanging-wall block. in the hanging-wall block of the brnf, the bend of the thrust-fault structures may be characterised as a roll-over anticline. note that the br06 hanging-wall flat transforms into a hanging-wall ramp (br06hwr), which was thrust-displaced along the upper footwall flat of the br04 thrust sheet (br04fwf). photograph: june 1984. 92 wall flat for about 200 m. as noted above, the br06hwf developed above the br05 thrust sheet at an early stage. the br06 thrust sheet is divided into two segments by the brnf (fig. 75). south of the brnf, the upper hanging-wall ramp of br06 was emplaced on the br04 upper footwall flat (fig. 75). the br07 thrust sheet was thrust piggyback onto the trailing-end segment of br05 and propagated up along the footwall ramp of br06. it has very chaotic internal structures dominated by polydiapirism. the position of the reference surface (l/r-unconformity) at an elevation of 20–30 m above sea level indicates that the br07 sheet was ramped up onto the flat above a duplex composed of the trailing segments of br03 and br05 (see later). there is only a thin cover of less than 5 m of the rubjerg knude formation on the back of br07, which is overlain by the br08 hanging-wall ramp. the br08 thrust sheet is the northernmost and uppermost sheet in the brede rende section. it is the smallest thrust sheet in the section, the thin frontal tip of the thrust sheet consists of the uppermost part fig. 76. ball-and-pillow structures superimposed by chaotic hydrodynamic brecciation in the upper part of the lønstrup klint formation in the brede rende section. the brecciation was formed by polysequential diapirism during thrusting of the br06 hangingwall flat (br06hwf ) over the hinge to the upper footwall flat of br05 (br05fwf) situated in the left side of the photograph. photograph: june 1993. 93 of the lønstrup klint formation, but with an up to 12 m thick succession of the rubjerg knude formation on top of the l/r-unconformity. due to the bend up along the br06 footwall ramp, the inclination of the br07 and br08 thrust sheets is c. 25°n. sedimentary units the lønstrup klint formation is strongly affected by ball-and-pillow load structures and hydrodynamic brecciation. the maximum thickness of the formation exposed is only about 20 m (tentatively measured in br02). the thickness of the rubjerg knude formation varies from thrust sheet to thrust sheet, indicating differential thrust-fault movement that either closed the piggyback sedimentation and/or lifted the formation up to a position exposed to erosion. the rubjerg knude formation was also subjected to hydrodynamic brecciation. at the top of the central part of the brede rende section, a glacitectonite and associated glaciotectonic imbrications are interpreted to be related to the advance of the norwegian ice; the sandy till is interpreted to be the kattegat till formation. lønstrup klint formation the lower and intermediate parts of the lønstrup klint formation are characterised by dark clayey mud. the interval 5 to 10 m below the l/r-unconformity is dominated by a few thick beds of light coloured sandy turbidites, and the uppermost 5 m is formed by thinbedded sand beds interbedded with mud. the size of fig. 77. small-scale ball-and-pillow structures distorted and intruded by water-escape injection. note that dish structures were formed above the water-escape pipe. detail of chaotic brecciation in the upper part of the lønstrup klint formation; frontal part of the br06 thrust sheet in the brede rende section. photograph: june 1998. fig. 78. small-scale disharmonic undulations formed by polysequential diapirism in the thinly interbedded clays, silts and finegrained sands of the upper part of the lønstrup klint formation. note the fold accentuation of the climbing ripple lamination in the central part of the figure. frontal part of the br06 thrust sheet in the brede rende section. photograph: june 1998. 94 ball-and-pillow structures is typically related to the initial thickness of the sand beds, and the subsequent hydrodynamic brecciation and chaotic structures formed during water-escape activities (figs 68, 76–78). the l/r-unconformity at the top of br04 truncates a large ball-and-pillow structure at the top of the lønstrup formation just north of the normal fault (fig. 79). this implies that some of the load structures, and possibly also initial water-escape dynamics, had commenced prior to the development of the unconformity. it may be that this phase of ball-and-pillow formation was initiated by the drainage of the large lake basin (see sadolin et al. 1997). thus the initiation of balland-pillow formation can be viewed as the consequence of vibration created by an increased water transport over the beds. at this locality, the formation of ball-and-pillow structures was clearly not the effect of loading by over-thrusting, but only the result of density variation of the primary sedimentary layers, since the top of the lønstrup klint formation was undergoing erosion and the gravel bed on the l/runconformity was deposited subsequently. rubjerg knude formation the rubjerg knude formation reaches a thickness of 15 m in the upper part of the br04 thrust sheet, but in the rest of the brede rende section it is less than 10 m thick. the relatively thin nature of the formation (3–5 m) in br02, br03, br05 and br07 is interpreted to indicate that these basins were over-thrust or thrustelevated at an early stage of thrust propagation. in contrast, deposition persisted in the piggyback basins of br04 and br06 before their upper footwall flats were overthrusted and deposition ceased in the basins. structures three types of structural features are described from the brede rende section: (1) diapir structures, including mesoscopic-scale sequential polydiapirs and hydrodynamic brecciation, (2) the brede rende normal fault (brnf), and (3) frost wedges. fig. 79. a large ball-and-pillow structure in the upper part of the lønstrup klint formation, truncated by the l/r-unconformity. this relationship demonstrates that at least part of the loading occurred prior to the thrust-fault emplacement. photograph: june 1998. facing page – upper: fig. 80. normal fault network in the footwall block of the brede rende normal fault developed in the br04 thrust sheet. photograph: june 1998. facing page – lower: fig. 81. frost wedges recognised in the rubjerg knude formation. the one on the right side of the spade (a) has well-developed, ‘upwards-fanning’ small-scale normal faults, whereas the one to the left of the spade (b) is a 5–10 cm wide fracture with a sand-fill. photograph: june 1997. 95 96 diapir structures the term diapir as used here follows the definition of weinberg & schmeling (1992 p. 425): “diapir is the non-genetic geological term applied to ductile intrusive structures. many diapirs may develop due to rise of gravitationally unstable buoyant fluids through denser overburden. such gravitationally unstable configurations consisting of viscous layers are known as rayleigh-taylor instabilities.” in the brede rende section, the diapir structures can be divided into two types: (1) smallto medium-scale diapirs that developed into hydrodynamic breccias in which primary sedimentary lamination is locally preserved, although distorted and irregularly folded, and (2) mediumto large-scale diapirs where mobilised mud intrudes overlying stratigraphic levels or thrust units. the first type of diapirism corresponds to the sequential polydiapirs of weinberg & schmeling (1992). these are initiated as small undulations or even flame structures, that develop into irregular upright folds with numerous minor undulations on their flanks (figs 76, 78). when the viscous mud broke through the bedding it formed intrusive pipes (fig. 77), and either spread out laterally between layers or released water, forming dish-and-pillar structures in the overlying beds fig. 82. the shift in tilts of bedding in the piggyback basin of the br08 thrust sheet is interpreted to reflect the propagation of ramps. the strike is the same (110°), but the dip of the lower sand beds is 40°, corresponding to deposition during propagation along a flat, whereas the dip of beds above the truncation surface is only 28°, corresponding to deposition during ramping. photograph: june 1997. fig. 83. back-thrust reverse faults displacing the sand beds in the rubjerg knude formation deposited in the piggyback basin of the br08 thrust sheet. these reverse faults are interpreted to have formed during the thrust propagation of the footwall ramp of br06/br05. photograph: june 1997. 97 (fig. 77). in the brede rende section, this type of diapirism occurs commonly in the upper sand-rich part of the lønstrup klint formation and in the rubjerg knude formation. a good example occurs at the tip of the br05 thrust sheet, where the thrust fault (br04fwf/br05hwr) is completely obscured by hydrodynamic brecciation. examples of the second type of diapirism include the brede rende diapir and the kramrende diapir. here the clay-rich units of the lower part of the lønstrup klint formation became mobilised by over-pressured water (or gas) to form an intrusive grey, homogeneous mud. the diapirism in the brede rende section was formed syntectonically during ramping (pedersen 1987). the thrusting displaced some of the feeders in the mushroom-shaped diapirs, and some of the diapirs intruded through the thrust sheets up into the thrust sheet above. moreover, the mushroom-shaped diapirs penetrate the l/r-unconformity at the top of the diapir (fig. 73). it may also be noted that some of the diapir feeders have been tilted by the bending produced by ramp propagation (fig. 4; pedersen 1987). brede rende normal fault the brede rende normal fault (brhf), in the central part of the brede rende section, is a planar fault that strikes 100° and dips 45°s (fig. 75). the vertical displacement is c. 20 m when measured from the hanging-wall (thrust-fault) flat of the br06 in the footwall block of the normal fault to the same flat in the hanging-wall block of the brnf. a network of smaller normal faults with minor displacements occurs in the footwall block (fig. 80) and adds to the monoclinal bend in the footwall block of the brnf. in the hangingwall block, the br06 thrust sheet is dragged along the fault plane and the drag is bounded by a minor splay fault. moreover, a weakly developed rollover-anticline outlined by the br06 thrust sheet occurs in the hanging-wall block of the brnf (fig. 75). above the northern limb-bend of the rollover-anticline, a minor depression (c. 5 m deep) was formed. in this depression, a series of minor imbricate sandy mud slumps formed, which may be viewed as synto epitectonic deposits at the top of the rubjerg knude formation in br06 related to faulting of the brnf. frost wedges frost wedges or fossil ice wedges are recognised in the rubjerg knude formation, as preserved in the upper part of the br04 thrust sheet. the clif f section here became exposed after the cross-section (plate 1) was drafted and is thus not included. it would have been situated near point 3975 m in the cross-section. the frost-wedge fractures are 5–10 cm wide and are filled with structureless sand. along the sides of the fractures, the bedding in the sand is bent down towards the fracture due to minor displacements along small fanning normal faults; the vertical range of the frost wedges is about 1–3 m (fig. 81). the presence of frost wedges in the rubjerg knude formation clearly indicates that the sand was ground frozen, and thus also elevated above water level in the glaciofluvial and glaciolacustrine environment that prevailed during the deposition of the formation. the ground-frozen condition of the sand may be the reason for the excellent preservation of the normal fault network related to the brnf. interpretation of structural development the first thrust sheets to move were probably br03 and br06, which ramped up to the upper footwall flat and moved southwards over a thin cover of the rubjerg knude formation. in the balanced cross-section, the presence of the long, thin br03 thrust sheet, and especially br06, requires that there has to be underlying lower and intermediate duplex sheets. the balanced cross-section model favours a continuation in the subsurface of several segments of the lower thrust duplex. the br03 thrust sheet is viewed as a coherent thrust sheet, which from the ramp of the minor br02 thrust sheet, continues along the lower décollement surface at the 30 m level. the lower trailing-end duplex segment extends northwards to the thrust fault separating the brede rende and the sandrende sections (sr01hwr/br08fwr). this trailing segment of the br03 thrust sheet is estimated to be about 300 m long, and the remaining five thrust sheets in the brede rende section have all been ramped up onto this segment along which the allochthonous transport and piggyback displacement took place. the simplest model for understanding the framework of the duplexes is to accept segmentation of the trailing end of the br05 thrust sheet. it is necessary that br06 was thrust over br05 before the trailing 98 end of br05 was thrust up over the footwall ramp of br04. the existence of the intermediate br04 hanging-wall ramp indicates that br04 had to ramp up two footwall ramps in dif ferent positions of the trailing end of br03. thus the model indicates that a lower hanging-wall ramp of br04 was emplaced along the intermediate footwall flat of br03. according to the construction of the balanced cross-section, this also necessitates a lower duplex segment to be thrust up in front of the lower hanging-wall ramp and flat of br04. these dif ferential thrust displacements provide an explanation for the development of the brnf. the displacement along the brnf is consequently considered to be due to two factors. the first 10 m offset was caused by normal faulting in front of br05, where a foreland-dipping bend of the br06 hanging-wall flat was created over the nose of the br05 thrust sheet. the next 10 m displacement was caused by a foreland-dipping limb of the tip of a duplex segment situated beneath the br04 thrust sheet causing the br04 hanging-wall flat to act as a normal fault. at the north end of the section, the br07 thrust sheet, which only has 3–4 m of the rubjerg knude formation at the top, was overthrust by the br08 thrust sheet at an early stage. the thickness of about 10 m of rubjerg knude formation on top of br08 indicates that after the two thrust sheets were thrust-separated, deposition of rubjerg knude formation continued in the piggyback basin of br08. this sedimentation probably took place while br08 in a piggyback position on br07 ramped over a lower footwall ramp of br03 and propagated along an intermediate flat, passing over the footwall ramp of br05/br06, before the temporary cessation of thrusting. in the br08 piggyback basin, the propagation of ramps is reflected in the change in tilt of the bedding (fig. 82). moreover, a number of minor back-thrusts have been recognised in these beds (fig. 83), and are considered to have been related to the ramp propagation. in the dynamic development of the brede rende section, both the frontal southern and the northern parts were involved in diapirism. in both parts, it is fig. 84. the sandrende diapir developed in the sr02 thrust sheet. the arrow indicates the direction of reverse faulting, which marks the prominent back thrust. along the steep northern flank, the hanging-wall ramp of sr03 (sr03hwr ) was bent. the bend of the l/r-unconformity formed due to the fold-bend-folding of sr02 at the lower footwall ramp hinge. 99 evident that the diapirism was active after the emplacement of the thrust sheets, since the hanging-wall flats are penetrated by diapirs rising from a mobilised underlying thrust sheet. however, it is also evident that the diapirism ceased before the maximum compression of thrust sheets had occurred. the termination of thrust compression was reached when the maximum inclination of the flats occurred. this coincided with the conclusive accumulated ramping of piggyback thrust sheets. thus, the inclined position of the feeders to the mushroom-shaped diapirs indicates a synthrust intrusive emplacement. it is therefore concluded that the diapirism was activated by ramp propagation and that some of the diapirs can be regarded as extreme developments of hanging-wall anticlines created during soft sedimentary deformation (fig. 74). sandrende section the sandrende section is one of the most studied parts of the lønstrup klint cliff section (fig. 5; houmarknielsen et al. 1996; sadolin et al. 1997). even so, the development of this section is not fully understood, and some new and revised details are added here. the main feature of the section is a broad basin containing a thick succession of the rubjerg knude formation deposited in a piggyback basin. to the south, a diapir distorts this basin, and to the north the basin is over-thrust by a thrust sheet of the stenstue rende section. the central part of the section preserves a remarkable development of normal faults. these were formerly regarded to have formed in response to the volume adjustments in the sandrende diapir (sadolin et al. 1997), but are now interpreted as elements of a thrust-fault propagation model with differential duplex segments ramping in the subsurface. tectonic architecture the sandrende section comprises four thrust sheets (sr01–sr04). the southern boundary of the section is the trailing-edge ramp of br07 and br08 in the brede rende section, and the northern boundary is the rather steep (> 60°) trailing-edge ramp of sr04. the boundary with the stenstue rende section to the north is a combination of this trailing-edge ramp and the hanging-wall flat of the frontal southernmost thrust sheet in the stenstue rende section (see below). the transition between the brede rende section and the sandrende section in the subsurface is not clear due to uncertain relationships between br07– br08 and sr01. the description below is based on the preferred interpretation, which traces the trailingedge ramp of br08 in the brede rende section down to the décollement surface 30 m below the reference surface. this implies that the lowermost trailing ends of br07 and br03 remain as low-lying segments that sr01 had to ramp over. an extra segment and some smaller adjustment splints of the sr01 thrust sheet were also left in the subsurface. this is reflected in some of the structural features exposed in the section between sr01 and sr02. at the tip of the sr01 thrust sheet, the lønstrup klint formation forms a thin wedge, which indicates that the initial hanging-wall ramp (sr01hwr) only had a dip of about 10°. however, after ramping was concluded, the thrust fault was steepened to the present dip of 40°n; the measured orientation of the ramp is 108°/40°n. the displacement of sr01hwr along the upper footwall ramp and flat of br08 is c. 53 m. the frontal part of sr01 has a bend, and it only dips about 25°n due to the change in thrust-fault inclination passing the upper footwall ramp hinge and the subsequent introduction of a small satellite thrust fault displacing the lower part of sr01 up over the tipwedge. the consequence of thrusting the thin c. 50 m long frontal part of the thrust sheet is that in the balanced cross-section, a lower duplex segment (sr01u) must be accounted for, and that sr01u at an advanced stage of sr01 thrust propagation was picked up in the thrust translation (see below). the sr02 thrust sheet was formerly interpreted as a large-scale diapir (sadolin et al. 1997; fig. 84). in the present structural analysis, sr02 is treated as one large thrust sheet in which the mobilised mud underwent mud diapirism at a relatively late stage. this assumption permits an approximation of balancing the thrust sheets, accepting that the thrust faulting is evidently the most important part of the dynamic development. the argument for this is based on the fact that sr02 over-thrust the back of sr01 with a displacement of about 100 m. this 100 m of displacement has to be compensated for by the same amount of displacement along the lower décollement surface, which can be calculated to have taken place at a stratigraphic depth of 30 m below the l/r-unconformity. the thrusting of sr02 resulted in a considerable amount of elevation during propagation along footwall ramps (sr01fwr and br03fwr), since the l/r reference surface is sit100 uated about 35–40 m above sea level in the cliff section. thus the ramping and displacement along the upper flat took place before the final emplacement of the lower duplex segment of sr01, indicated by the normal fault displacement of both sr01 and the frontal part of sr02. as noted in the kramrende section description, steep ramping creates back-thrusting at the hinge of the hinterland-dipping limb. thus the peculiar mushroom-shaped structure with a wing pointing to the north is considered to be the effect of reverse faulting due to back-thrusting (fig. 84). the reverse fault feature may have been accentuated by re-orientated internal detachment folding and irregular diapirism in the sandrende diapir. furthermore, it should be noted that the l/r-unconformity surface has a steep dip on the northern flank of the sandrende diapir. near the beach level, the l/r-unconformity bends into a gentle dip indicating that in the trailing end of sr02, the lower hanging-wall flat rests on the lower footwall flat coinciding with the décollement level at 30 m. the sr03 thrust sheet is relatively small with a displacement of about 75 m. the tip of the thrust sheet is bent upwards into a nearly vertical position due to drag along the almost vertical northern flank of the sandrende diapir (sr02). thus the sr03 thrusting was rather early, but as the rubjerg knude formation is about 10 m thick in sr02 there was a significant time span before sr03 was thrust up on the back of sr02. sr03 was displaced up along the upper footwall ramp, which is exposed in the cliff section. sr03 was also displaced along an intermediate flat situated at the 20 m level, indicated by the thickness of the thrust wedge. during thrust propagation, the trailing end of sr03 was cut off and left as an isolated duplex segment, while the frontal part of sr03 was displaced along the intermediate flat (see plate 2). sr03 was over-thrust by sr04 with a relatively short time gap, as indicated by the thin (3 m) succession of rubjerg knude formation on top of sr03. the thrust displacement of sr04 over sr03 is about 60 m, and fig. 85. conjugate normal faults developed in the lønstrup klint formation in the sr04 thrust sheet. an offset of about 1 m can be recognised by correlating turbidite sand beds in the footwall block to the same beds in the hanging-wall block. the normal fault framework is interpreted to be due to lateral extension in the sr04 thrust sheet during its propagation over the upper footwall hinge of an underlying duplex. photograph: may 1995; measuring staf f divisions (centre) are 20 cm. 101 the accumulated displacement of the trailing end of sr04 relative to sr02 is in the order of 135 m. the frontal part of sr04 consists of a relatively thin wedge of the upper part of the lønstrup klint formation overlain by an up to 28 m thick succession of the rubjerg knude formation. in the central and rear parts of sr04, the thickness of the lønstrup klint formation increases to more than 20 m, indicating the existence of a hanging-wall ramp which can be traced down to the décollement zone, 30 m below the l/r reference surface. the central part of sr04 forms a broad hanging-wall anticline, where a number of extensional normal faults cross-cut the lønstrup klint formation (fig. 85). the southernmost normal fault in this system is considered to reflect the foreland-dipping features formed due to displacement of the hangingwall anticline along the intermediate flat. finally, it should be noted that the piggyback basin (rubjerg knude formation) of sr04 is divided into two subbasins, one in the southern frontal part and one in the northern trailing part of the thrust sheet. the area between the sub-basins lacks the rubjerg knude formation because it corresponds to the crest of the hanging-wall anticline. sedimentary units the type sections of the lønstrup klint and rubjerg knude formations, as defined in this bulletin and previously described by sadolin et al. (1997), are situated at sandrende. as defined above, this succession is divided here into the lønstrup klint formation and the overlying rubjerg knude formation, which are separated by the l/r-unconformity (fig. 19). the rubjerg knude formation is covered by an up to 1 m thick homogeneous sandy till, which is referred to the kattegat till formation (fig. 30). lønstrup klint formation in the sandrende section, the lower exposed part of the lønstrup klint formation is composed of laminated clayey to sandy mud, intercalated with a few thin sandy turbidites that grade up into finely laminated clay-rich mud. in the upper part of the formation, thicker turbidite sand beds with climbing ripples give the formation a banded light/dark coloured appearance (figs 19, 85). only very few load structures and hydrodynamic breccias have been noted in the sandrende section, except in the lower part of the sr04 sheet where ball-and-pillow structures and small-scale polydiapirism have been observed (fig. 86). the ball-andpillow features are about 20 cm in thickness, which is probably the thickness of the original beds; they are typically elongated about 50–75 cm parallel to the strike of the bedding, suggesting they were formed during the thrust deformation. rubjerg knude formation the rubjerg knude formation comprises three units: (1) a lower unit c. 5 m thick consisting of trough crossbedded sand and gravel, (2) a middle unit dominated by climbing ripple cross-laminated sand, and (3) an upper unit comprising alternating beds of small-scale ripple cross-laminated sand and trough cross-bedded sand (fig. 19). the units reflect the change from fluvial to lacustrine and back to fluvial depositional environments (sadolin et al. 1997). the rubjerg knude formation has an onlapping relationship in the frontal part of the sr04 thrust sheet, which reflects initial thrust faulting during sedimentation (sadolin et al.1997). in the central part of the sr04 thrust sheet, growth-fault sedimentation along normal faults is recorded in the lower part of the rubjerg knude formation. the growth faults coincide with the foreland-dipping limb of the hanging-wall anticline of sr04 (fig. 87). this syntectonic sedimentation supports the piggyback basin concept for deposition of the rubjerg knude formation. moreover, a slumped block 0.5 × 2 m in size occurs along one of the normal faults indicating that the tip of the satellite thrust in sr04 was exposed to erosion and slumped into the basin. similar slumped blocks were observed on the northern flank of the sandrende diapir indicating that the diapir rose above the depositional surface during emplacement and that fragments of the lønstrup klint formation slumped into the piggyback basin. the synsedimentary rise of the vertical diapir wall was also reflected in sedimentation of small point-bar wedges along the vertical flank of the sandrende diapir. towards the top of the rubjerg knude formation, broad trough cross-bedding is observed. minor thrust faults splaying out from the tip of ss01 displace the cross-bedded sand, and the base of some of the troughs dramatically truncate the thrust faults, in a similar fashion to that observed in the br04 piggyback basin of brede rende (see above). at the top of the rubjerg knude formation, sand was deposited in a 102 fig. 86. mobilisation and small-scale polydiapiric features developed in the upper part of the lønstrup klint formation in the sandrende section (trailing end of sr04). the polydiapirs started along a bed of clayey mud as small flames (with small wavelength), which were subsequently folded around the taller diapirs. the sandy beds above and below constitute planar laminated and climbing ripple cross-laminated fine-grained sand with organic debris and mud draping the ripples. locally in this sand, hydrodynamic mobilisation has created zones of mud-free structureless sand and the accumulation of mud forming dendritic structures. the dynamic development of the structure is illustrated in fig. 88. fig. 87. extensional normal faults with related growth-fault sedimentation of sand and gravel in the lower part of the rubjerg knude formation. the growth faults are marked with arrows indicating the direction of displacement. the top of the rubjerg knude formation in the sr04 thrust sheet is over thrusted by the ss01 thrust sheet. the two sheets are separated by a thrust fault that acts as footwall flat of sr04 (sr04fwf) and hanging-wall ramp of ss01 (ss01hwr). photograph: may 1995. 103 depression above the top of the sandrende diapir. this depression was probably formed by relaxation collapse of the diapir during consolidation and dehydration. structures and breccias structural investigations in the sandrende section focused mainly on the normal faults and their relationship to the thrusting, diapirism in the sandrende diapir, smallscale incipient polydiapirism, and the record of a deep frost wedge cutting the rubjerg knude formation. normal faults the sandrende section is an important locality for the investigation of normal faults formed on the forelanddipping limb of hanging-wall anticlines. thus, one set of normal faults displaces the frontal parts of sr01 and sr02, and another set displaces the central part of the sr04 thrust sheet. in the frontal part of sr01, a hanging-wall anticline developed due to ramping from the 10 m to the 20 m flat level. the normal faults here displace the rubjerg knude formation of sr01 as well as the tip of sr02. the faults now have a dip of about 45°s, but initially probably had a much steeper dip (up to 80°) subsequently reduced during the final bend of the br08 footwall flat. in the trailing end of sr01, a steep normal fault displaced sr01, as well as the over-thrust sr02, with an offset of 10 m. this probably reflects the influence of a sub-surface duplex, similar to the development of the brnf. the normal faults in the sr04 thrust sheet can be fig. 88. the polydiapiric structures and hydrodynamic breccias shown in fig. 86 are interpreted to have developed in the following five steps. 1: initial sedimentation of a clayey mud bed in a succession of mud and fine-grained sands. 2: first-order formation of small flames can be regarded as micro-diapirs with small wavelength. 3: second-order small diapirs developed with increased wavelength. small-scale thrusting and overturned geometry indicates formation during thrust-fault propagation. 4: increased mobilisation creates small-scale domes with extensional fractures forming in the crest. 5: liquefaction of the heterolithic sediment results in segregation of the sand and mud components. the mud accumulates in an irregular diapir from the top of which the mud-saturated liquid intrudes laterally along the primary parallel lamination. some mud and fragments of sand fall to the base of the diapir under gravity. 104 viewed as two sets of a fault framework. the first set formed 45–60°s dipping faults with displacements of 1–3 m. the southerly dipping tilt of the l/r-unconformity is regarded as the foreland-dipping limb of the hanging-wall anticline formed in sr04 (fig. 87). the faults above the foreland-dipping surface developed as growth faults associated with syntectonic sedimentation, as recorded in the lower part of the rubjerg knude formation. during displacement along the normal faults, a minor satellite thrust cross-cut the sr04 thrust sheet, and the tip of the satellite thrust sheet was slump-faulted to form slumped blocks in the growth-fault setting of the piggyback basin. the normal fault network at the crest of sr04 is very impressive (fig. 85). the strike of the normal faults is 090° with a dominant dip of 50°s, although a small number of conjugate faults with a dip of 60– 75°n also occur. in view of the angle of conjugate faulting, these normal faults could have formed due to the loading of the ss01 thrust sheet emplaced above the upper footwall flat of sr04, but could also have formed due to necessary extensional adjustments during propagation over the hinge of the footwall ramp. diapir structures the sandrende diapir only af fected one thrust sheet (sr02), in contrast to the brede rende and kramrende diapirs where two or more thrust sheets were involved in the diapir formation. the most impressive feature of the sandrende diapir is the major back-thrust, which has an offset of about 20 m towards the north. initially it was probably an almost vertical reverse fault, which was re-orientated and accentuated during thrustfault propagation. a number of smaller reverse faults occur along the steep northern wall of the diapir, which internally is composed of mobilised mud. in the upper part of the diapir, distorted bedding-structures isolated as ‘xenoliths’ in the upper part of the diapir are interpreted as fragments of hanging-wall anticlines. judging from the thickness of the diapir feeder, the diapir formed over a hanging-wall ramp where the sr02 thrusting ramped from the upper 10 m flat to the lower 20–30 m flat level. the formation of diapirs was evidently initiated by mobilisation on a small scale. an illustrative small-scale example of diapirism was observed in the northern part of the sr04 thrust sheet (fig. 86) where mobilisation and small-scale polydiapirs developed in the upper part of the lønstrup klint formation in the sandrende section (trailing end of sr04). the polydiapirs are related to beds of clayey mud deposited between the thicker beds of sandy turbidites. along the boundary of the 25–75 cm high diapirs, small flame structures occur and the laminated sandy beds above are irregularly folded. locally, hydrodynamic mobilisation created mud-free structureless sand and complex mud structures developed. an interpretation of the dynamic development of the structures is given in fig. 88. frost wedge a 20 m deep fracture cross-cuts the rubjerg knude formation in the central part of the southern sub-basin in the sr04 thrust sheet. the fracture is less than 10 cm across, and can be followed as an irregular trace downwards into the sand sequence with a number of minor lateral jumps. this irregular fracture is one of the few structures that can be interpreted as a frost wedge. it does not penetrate the overlying ss01 thrust sheet, indicating that it formed within the rubjerg knude formation from an exposed surface downwards into a freshly frozen sand package. it can be inferred that during the latest phase of thrusting, the sr04 thrust sheet was elevated to a position such that the top of the rubjerg knude formation was exposed subaerially. interpretation of structural development a hanging-wall anticline developed c. 40 m from the tip of sr01 when it passed the footwall ramp and flat of br08. this initially created a foreland-dipping tilt of the sr01 thrust structures, and was also responsible for the formation of the normal faults described above. however, the frontal part of sr01 has to be accommodated with a duplex segment in the subsurface (sr01u). the trailing end of sr01 is rooted down to the décollement level, where it corresponds to the segment adjusting the c. 80 m long frontal part of sr02. when the hanging-wall ramp of sr02 initiated the propagation up along the footwall ramp, a hangingwall anticline was formed that developed into the sandrende diapir with its marked back-thrust. during the sr02 propagation along the footwall ramp, the sr01u-duplex was pushed up and created a minor hanging-wall anticline, along which foreland-dipping limb a normal fault developed and displaced the sr02 thrust sheet as well as sediments in the piggyback basin of sr01. 105 the lower 10 m of the rubjerg knude formation was deposited throughout the sandrende section, with the exception of the northern part of sr03, which had already been blocked by thrusting of sr04. displacement of the lower hanging-wall ramp of sr04 onto the intermediate flat of sr03 (and sr02) then took place. subsequently, the first normal growth faulting was initiated at the margin of the southern part of the piggyback basin above the foreland-dipping l/r-unconformity. propagation of the lower sr04 hanging-wall ramp separated the piggyback basin into two sub-basins where deposition of the upper part of the rubjerg knude formation took place, while the crest of the anticline between the sub-basins was probably subjected to erosion. the southernmost thrust of the stenstue rende section (ss01) over-thrust the top surface of the rubjerg knude formation (sr04fwf) as well as the eroded surface of the ramp anticline; this prevented deposition in the sr04 piggyback basin. ramping of the lower trailing segment of sr03 was activated in the latest stage of thrusting. the propagation of this duplex segment (sr03u) for a short distance up along the footwall ramp contributed to the flat-topped hanging-wall anticline formed in sr04. this final duplex emplacement may have been one of the causes for the formation of the normal fault framework in the hanging-wall anticline of sr04 (fig. 85). stenstue rende section in the stenstue rende section, a remarkable and dramatic episode of megaslumping is recorded. formation of a very large southward-verging anticline was accompanied by chaotic hydrodynamic brecciation (fig. 89). another important element related to this section is the c. 200 m displacement of the frontal thrust sheet over the sandrende section to the south. fig. 89. the large slump fold in the stenstue rende section. the slumping folded the ss05 thrust sheet into an overturned anticline during displacement down the normal fault escarpment. the escarpment was formed during normal fault displacement of the tip of ss04 parallel to the foreland-dipping limb of a hanging-wall anticline in ss03 (see plate 2 and fig. 90). photograph: june 1999. 106 the stenstue rende section is named after the gully situated between the stenstue rende section and the sandrende section leading inland from the beach. in the northern part of the section is the gully known as the søndre grønne rende. this is reached by a path through the pinewood connecting with the main road between rubjerg and lønstrup. tectonic architecture the stenstue rende section comprises six thrust sheets (ss01–ss06). to the south, the footwall ramp and flat of sr04 in the sandrende section bound the section. to the north, the boundary is defined by the footwall ramp of ss06, which coincides with the hanging-wall flat of the southernmost thrust in the grønne rende section (gr01). the most important thrust sheet in the stenstue rende section is the more than 400 m long ss01 thrust sheet, the frontal part of which over-thrust the northern part of the sandrende section and has a displacement of more than 200 m. the initial ramping of the ss01 thrust sheet was located at a gently dipping hanging-wall ramp. subsequent to the ramping, part of the tip was eroded away during the uplift exposure of the hanging-wall anticline above the ramp and the final truncation of the glaciotectonic unconformity. due to the increase in thickness of the ss01 thrust sheet, corresponding to a change from the 10 m upper flat level to the 20 m flat level, an intermediate hanging-wall ramp developed about 100 m from the frontal tip. this hanging-wall ramp rests on top of the footwall flat (sr04fwf) above the prominent normal fault structure in the sandrende section. only a small remnant of the northern part of the upper flat structure is preserved, and this is not very well exposed due to its location in the inner part of the stenstue rende. the lower ss01 hanging-wall ramp (ss01hwr, ramping from the 30 to 20 m flat level) situated in the middle part of the ss01 thrust sheet is now exposed in a steeply dipping position along the sr04 footwall ramp. the propagation of ss01hwr was responsible for the bend of the footwall syncline in the rubjerg knude formation in sr04, and the subsequent tilting of ss01hwr resulted in the appearance of a more or less vertical boundary between the two sections. the vertical orientation is a combination of 45° dip on the footwall ramp added to 45° dip on the hanging-wall ramp. note that in the balanced cross-section, there is a c. 200 m long lower ss01 duplex segment (ss01u) which needs to be allowed for. this implies that after ramp propagation, the lower hanging-wall flat of ss01 was displaced along the intermediate footwall flat on ss01u. ss02 is a small thrust sheet, thrust onto the footwall ramp of ss01; this footwall is composed of the rubjerg knude formation situated in the upper part of ss01. note that the frontal part of ss02 has a surprising vertical orientation and is displaced by a more or less horizontal extensional fault, the cause of which is discussed below. the frontal part of the ss03 thrust sheet is shown in the cross-section as a rather simple, upright thrust structure (plate 1). however, the ss03 thrust sheet is in reality a chaotic load and hydrodynamic breccia complex. at the base of the cliff section is an upright anticline, which is regarded as a key structure in the interpretation of the thrust development (fig. 90). above the anticline, a normal fault dipping 40°s truncates the c. 30 m thick rubjerg knude formation. the thin frontal part of the ss04 thrust sheet is characterised by chaotic brecciation. the trailing end is c. 10 m thick, dipping 45°n, with the hanging-wall flat thrust along the footwall flat of ss03. the tip is separated from the trailing part of the ss04 thrust sheet by a 40° dipping normal fault with a displacement of about 50 m. this normal fault formed an escarpment truncating the rubjerg knude formation on top of the ss03 thrust sheet, upon which deposition of a coarse clastic breccia took place (fig. 91). the normal fault escarpment was finally overridden by the frontal part of the ss05 thrust sheet, which slump-thrusted down the fault plane and formed a major overturned slump fold (fig. 89). the formation of the megaslump fold took place after the ss05 thrust sheet was thrust up along the footwall flat of ss04 to the head of the escarpment from where it gravitationally slid down to the depression on the back of the ss04 tip. a soft sedimentary tectonic breccia was formed at the transition between ss04 and ss05, which was cross-cut by a number of minor steeply southward dipping normal faults reflecting the final settling of the fault-slump structure. the ss06 thrust sheet is about 30 m thick, its lower hanging-wall flat resting on the footwall ramp of ss05. it has a steep dip and has been strongly disturbed by mobilisation and internal diapirism. this thrust sheet is included in the stenstue rende section because it involves the trailing lower duplex segments of ss05 and ss04. from the position of the l/r-unconformity surface, about 30 m above sea level, it can be inferred that the ss06 thrust sheet was raised up over the low107 er trailing segments during ramping and subsequent stacking of a subsurface duplex complex. sedimentary units the most interesting sedimentological feature within the stenstue rende section is the record of syntectonic sedimentation related to normal faulting. this includes slump deposits as well as a gravel bed developed on the escarpment surface of the normal fault. as these sedimentary features are related to deposition in the piggyback basin, they are described below as part of the rubjerg knude formation. the sediments of the lower lønstrup klint formation have been strongly affected by thrust shearing, and the upper levels were modified by hydrodynamic brecciation. the rubjerg knude formation comprises a confusing mixture of redeposited units together with the main fluvial-lacustrine sediments related to the piggyback basins. lønstrup klint formation the lower part of the lønstrup klint formation is exposed in the ss06 thrust sheet, where the lower hanging-wall flat is thrust up along the footwall ramp of ss05. here bluish grey clay alternates with dark redbrown clay in a laminated to thin-bedded unit (fig. fig. 90. the crest of the hanging-wall anticline formed in the ss03 thrust sheet in the stenstue rende section. photograph: june 1984; the staf f divisions are 20 cm. fig. 91. the conglomerate/breccia formed along the fault escarpment truncating the ss03 thrust sheet. photograph: june 1984; the staff divisions are 20 cm. 108 92). it is evident from the shear structures that this unit acted as a décollement zone during thrusting. the main part of the lønstrup klint formation exposed in this section comprises the upper sand-dominated part of the succession. the breccias in the ss03 and ss04 thrust sheets probably initially formed as mediumto large-scale ball-and-pillow structures in sand beds 20–60 cm thick during initial thrusting; the formation was subsequently deformed during gravity slumping. rubjerg knude formation nearly 20 m of fluvial-lacustrine sand were deposited in the piggyback basin of ss01 and ss03 during the thrust-fault activity af fecting the stenstue rende section. however, the most conspicuous unit is the remarkable conglomerate/breccia related to the normal fault. the gravel bed draping the fault escarpment is 10–50 cm thick and includes clasts up to 10 cm in size. locally, the clasts occur in a clayey mud matrix, but the latter has often been removed by recent erosion. it is perhaps surprising that a coarse gravel bed could have accumulated and been preserved along a fault escarpment dipping at about 35° (fig. 91). one explanation may be that the escarpment was only exposed for a very short time before the ss04 thrust sheet was displaced down the fault plane; in this case, the redeposited gravel rather represents a tectonic breccia composed of the smeared-out lithologies of the l/r-unconformity and surrounding sediments. the breccia is thus interpreted as the residue of a brecciated thrust sheet. the source of the clasts was probably the l/r-unconformity, and some of the material may have been derived from the unconformity by successive erosion during exposure at the head of the fault escarpment. this probably only occurred for a brief period before the escarpment was covered by the slump-slide of the ss05 thrust sheet. the small piggyback basin on top of the ss05 thrust sheet is a double syntectonic basin which was partly carried on the back of a thrust sheet as well as acting as a depression in the hanging wall of a normal fault. a 9 m thick succession represents the fill of this basin. the lowermost 3 m consist of large-scale cross-bedded medium-grained sand, rich in clay and silty mud clasts. towards the upper part of this unit, clay drapes on the cross-bed foresets become more common and the beds are affected by small-scale slumping. the overlying 5 m thick unit comprises sand beds 30–50 cm thick, with mud intercalations 5–20 cm in thickness. clay clasts are common and the sand shows smallscale ripples. the uppermost 1 m thick bed consists of mainly horizontal laminated sand and mud. this piggyback basin succession is interpreted to record a fluvial depositional environment that with time developed into a small shallow lake. a number of small south-dipping normal faults intersect the rubjerg knude formation up to the base of the thinly bedded muds and sands, indicating that the lake first became established when the fault activity ceased. structures four types of structures in the stenstue rende section deserve particular mention: (1) mesoscopic thrust-fault structures above the lower hanging-wall flat, (2) hanging-wall anticlines, notably the one in the central part of the section, (3) normal faults, the most important being the major escarpment-producing fault, and (4) slump folding related to the escarpment of the same fault. thrust-fault structures thrust faulting related to the décollement zone in the stenstue rende section has been observed in the lower hanging-wall flat of the ss06 thrust sheet. here the décollement zone is located in the 30 m flat level, which corresponds to the base of the 30 m thick lønstrup klint formation where lithologies are mud-dominated, comparing dark blue-green-greyish, clayey or silty mud with a few light grey coloured, fine-grained sand laminae. two types of structures are distinguished: imbricate duplexes and listric imbricate fans (figs 92, 93). the duplex imbricates appear within a 1 m thick unit bounded by thrust-shear surfaces below and above (fig. 92). the mesoscopic-scale duplex complex consists of sheets about 0.5 m thick and 1–3 m long. some of the duplexes are folded into antiformal stacks and form lensoid networks. the basal and roofing thrust faults occur as 20 cm thick shear bands penetrated by flat anastomosing jointing (fig. 92). the listric fans are outlined by 1–2 cm thick sedimentary layers or tectonically induced sand streaks (fig. 93). they rise from a narrow thrust plane, recognisable as a joint surface draped by a 1 mm thick film of black mud, and extend upwards into the muddy lithology where they seem to disappear before being over-thrust 109 by the next thrust joint surface about 1 m above the basal thrust surface. hanging-wall anticlines three hanging-wall anticlines have been recognised; the anticline in the frontal part of ss01 has been commented on above. the second example is not very obvious, but was developed above the intermediate hanging-wall ramp of ss01. the structures related to it were later modified by re-orientation due to the bend of ss01 up along the sr04 footwall ramp. the third hanging-wall anticline is the key structure in the stenstue rende section and is situated in the middle part of the ss03 thrust sheet. the ss03 hanging-wall anticline (fig. 90) was folded due to the ramping in the middle of the lower trailing duplex segment (ss01u). this ramping took place at a mature stage of thrusting, and ss01u was separated into two segments. the anticline is upright and tight, and the onlapping sedimentation of the rubjerg knude formation on the northern flank indicates that the ss03 thrust sheet had commenced transport along a footwall ramp and flat prior to the anticlinal folding. fig. 92. along the lower hanging-wall flat of the ss06 thrust sheet, an imbricate duplex complex has been recognised; bounding thrusts indicated by shear arrows . the thrust-fault imbrication formed in the lowermost part of the lønstrup klint formation during displacement along the décollement surface. the trowel is c. 30 cm long. photograph: june 1997. fig. 93. an imbricate fan formed in the lower part of the lønstrup klint formation in the ss06 thrust sheet. the trowel is c. 30 cm long. photograph: june 1997. 110 normal faults two normal faults are discussed: (1) the extensional fault with horizontal fault plane that displaces the tip of ss02, and (2) the major normal fault displacing the ss04 thrust sheet. the extensional fault affecting ss02 was formed north of the hanging-wall anticline developed over the intermediate hanging-wall ramp of ss01. it is interpreted to have formed initially as a normal fault dipping c. 45°n on the foreland-dipping limb of the ss01 hanging-wall anticline. subsequent to displacement on the normal fault, the ss02 thrust sheet and the fault were tilted into vertical and horizontal positions, respectively, during the fault-bend folding resulting from the ss01 propagation up along the footwall ramp. the major normal fault displacing ss04 is also regarded as a fault that developed on the foreland-dipping limb, here related to the anticline in ss03. it is observed that the l/r-unconformity dips beneath the beach level, indicating that the reference surface is not elevated and consequently that the underlying ss03 hanging-wall flat rests on a footwall flat; the normal fault is thus preserved with its initial orientation. the formation of the normal fault is similar to the formation of the brnf in the brede rende section (see above). slump folding the large-scale slump fold formed by the ss05 thrust sheet as it was displaced down the foreland-dipping fault escarpment can be compared to the same type of deformation described from the martørv bakker section. however, in the stenstue rende section, the lønstrup klint formation is still preserved as a coherent sheet, deformed into a major southerly overturned fold with an amplitude of about 15 m and an irregular fold axis orientated se–nw (c. 150°) (fig. 89). during slumping along the escarpment, the redeposited units were strongly affected by hydrodynamic brecciation resulting in the chaotic disorganised nature of the sediments. interpretation of structural development the important question in the development of the stenstue rende section is the time of formation of the ramp bend anticline during thrust-fault propagation. the interpretation given here is based on the description above, and the balanced cross-section and model for ramping in the subsurface given in the ramp cross-section (plate 2b). firstly, it should be remembered that there is evidence of a long translation along the décollement zone, primarily indicated by the considerable distance of ss01 transport over the upper footwall flat in the sandrende section (sr04). secondly, the displacement of the frontal part of ss01 must be balanced with a lower duplex segment (ss01u) in the subsurface. moreover, the displacement of ss01 also affected the ss02 thrust sheet by superimposed structural development. the superimposed model here advocated is supported by the following interpretation. as the initial angle of thrust faulting rarely exceeds 30° (jaeger & cook 1979), superimposed rotation must have affected the ss02 thrust. the angle between the ss02 hanging-wall ramp and the l/r-unconformity surface is about 30° indicating a normal type of thrusting when the rubjerg knude formation was horizontal. considering the thrust in this pre-rotated position, it is easy to envisage that the extensional fault offsetting the tip of ss02 as a normal fault formed over the lower hanging-wall ramp of ss01. to restore the thrust sheet into an upright position, two phases of rotation are necessary. the first one would be the ss01 ramping on the footwall ramp of sr04, and the second would be the re-orientation of the ramp due to the fault-bend provided by the thrusting of a subsurface segment of sr03 up along its footwall ramp in the sandrende section. the initial ramping of the leading edge of ss01 probably took place during sedimentation in the lower part of the piggyback basin of sr04. thrust propagation of ss02 must have been initiated at the same time, indicating that the ss03 thrust sheet in the trailing end of ss02 also participated in the translation along the 10 m flat level (ss01 intermediate footwall flat). during the translation of the lower footwall ramp in the trailing end of ss01, the 200 m long lower ss01u segment must also have been thrust, which is interpreted to have caused the ramping in the central part of ss01u. above this ramp, a lower hanging-wall anticline developed, which also folded the overlying ss03 thrust sheet into the exposed anticline in the middle part of ss03, resulting in the foreland-dipping footwall flat of ss03 and the initiation of normal faulting. part of the ss04 thrust sheet had by then already propagated over the ss03 footwall flat, and was therefore subsequently displaced by the normal fault with a drag down the fault plane. the displacement on the ss01u footwall ramp must have been relatively small to create and preserve an upright, close to tight anticline. if the dis111 placement had continued, it is likely that a more flattopped anticline would have developed and minor normal fault imbricates would have been the result, rather than the marked fault escarpment that actually formed on the southern flank of the anticline. the slump-thrusting of ss05 down into the depression on the hanging-wall block of the normal fault is interpreted to have taken place shortly after the ss04 was down-faulted. propagation of the ss05 thrust sheet was combined with the push on its footwall ramp by thrusting of the ss06 sheet. during this final thrusting in the stenstue rende section, the trailing lower segments were stacked into a duplex, probably analogous to the mesoscopic-scale duplex structure exposed along the hanging-wall flat of ss06 (fig. 92). grønne rende section in the geological cross-section of lønstrup klint presented by jessen (1931), two gullies were indicated south of the rubjerg knude fyr (the lighthouse), namely søndre and nørre grønne rende. by the year 2000, the cliff profile had no obvious gullies that these names can be attached to, although søndre grønne rende must have been close to the gully so annotated in the northern part of the stenstue rende section. the name grønne rende section is therefore adopted here to cover the section between stenstue rende and rubjerg knude fyr. the section comprises twelve nearly vertically orientated thrust sheets, which can be characterised as a listric imbricate fan. only the frontal parts of the thrust sheets are ramped up into steeply dipping positions, and in these parts of the thrust sheets the lønstrup klint formation is thin, whereas the rubjerg knude formation is relatively thick. the general impression of the section is of thin mud sheets alternating with thick units of sand (fig. 94). the points of interest in this section are the means of formation of an imbricate fan of uniform thrust sheets, and the mechanism by which the sheets reached their vertical orientation. also of interest is the arrangement of the now concealed duplex segments in the subsurface, where balancing of the thrust sheets indicates shortening of about 60%. fig. 94. view along the grønne rende section to the north where the lønstrup klint formation forms thin mud sheets interleaved with thick sand sheets referred to the rubjerg knude formation. in the far distance, stortorn forms the vertical cliff facing the sea. height of cliff is c. 50 m. photograph: august 1984. 112 tectonic architecture the grønne rende section comprises a leading-edge thrust sheet (gr01), which consists of a 30 m thick section of the lønstrup klint formation, succeeded above the l/r-unconformity by about 15 m of the rubjerg knude formation. north of gr01, a further twelve thrust sheets (annotated gr02–gr13) are exposed, each composed of an average thickness of c. 10 m of the lønstrup klint formation overlain by about 25 m of the rubjerg knude formation. to the south, the section is bounded by the thrust fault that separates the lower hanging-wall flat of the grønne rende frontal thrust sheet (gr01) from the footwall rampand-flat of the northernmost thrust sheet (ss06) in the stenstue rende section. to the north, the boundary of the grønne rende section is defined by the thrust fault that acts both as the hanging-wall ramp of rf01, the frontal thrust sheet in the rubjerg fyr section, and as the footwall ramp-and-flat of gr13. in the description of the thrust sheets, it is assumed that the thrust sheets initially involved only the lønstrup klint formation, and that the rubjerg knude formation was deposited syntectonically and separated into small piggyback basins between the sheets. as the imbricate thrust sheets constitute the most important structural element in this section, each thrust sheet is described separately in the structural account below. sedimentary units in the grønne rende section, the lønstrup klint formation is mainly represented by the upper levels of the formation. the only exception is the southernmost thrust sheet gr01, in which lower stratigraphic levels of the formation are also exposed. in this section, intra-rubjerg knude formation erosional surfaces locally incise the unconformity defining the lønstrup klint formation – rubjerg knude formation boundary. this unconformity is thus composite in places but the term l/r-unconformity is retained as it clearly still forms the boundary between these two formations. during hanging-wall ramping of the thrust sheet tips, the gravel beds on the unconformity were partly removed from the unconformity surface, and desiccation cracks may be present (only observed in the uppermost tips of the thrust sheets) indicating that some of the tips were exposed above water level. the rubjerg knude formation mainly comprises the same three units described in the sandrende section. however, a number of variations in sedimentary architecture occur due to syntectonic sedimentation. lønstrup klint formation the lower part of the formation exposed along the hanging-wall flat of gr01 consists of dark grey laminated mud with a few c. 0.5 m thick white sand turbidites; these have been strongly disturbed by thrusting, contortion and mud-mobilisation. the upper part of the formation is dominated by light-coloured sandy turbidites up to 1 m thick, interbedded with 10 cm layers of blue-grey clayey mud. rubjerg knude formation it has already been noted that the rubjerg knude formation was deposited in a number of small piggyback sub-basins. in the description of the piggyback basin architecture, four depositional elements are differentiated. 1. flat-parallel bedding (f-bedding): initially horizontal stratification of a bed deposited on a surface parallel to a flat as well as to the mean level of the l/r-unconformity. 2. ramp onlap (r-onlap): horizontal stratification or large-scale cross-bedding in a bed deposited on an inclined unconformity surface that had been tilted due to ramping prior to sedimentation. 3. foreland-dipping onlap (d-onlap): initially horizontal stratification in a bed deposited on an inclined unconformity surface dipping towards the foreland due to the repositioning of a hangingwall ramp on a footwall flat. 4. climbing ripple stratification (c-bedding): climbing ripple cross-lamination in beds 20–80 cm thick, commonly limited by f-bedding below and above (figs 24, 95). the rubjerg knude formation is interpreted as a glaciolacustrine deposit. sediment influx was probably relatively constant, and the sedimentary structures developed in the individual thrust sheets were governed by local conditions. during thrusting, the sedimentary base level changed, and the accommodation space varied depending on the size of the piggyback basins. thus the flow regime fluctuated and a variety of sed113 imentary structures formed, which are interpreted to reflect the thrust-fault development. the syntectonic banana-shaped basin described by pedersen (1987; fig. 4) was based on observations in the upper part of these piggyback sub-basins. these structures might also be characterised as footwall synclines that developed as growth-fault synclines, where deposition took place as the hanging-wall block was thrusted up along the footwall ramp dragging the underlying limb up along the thrust fault during the displacement. structures a systematic description of each thrust sheet in the section is provided below, together with some references to the syntectonic sedimentation. in general, the thrust sheets constitute an upright hanging-wall ramp, which initially had a dip of less than 20°. hydrodynamic brecciation and mud mobilisation occurred along the hanging-wall ramps and flats. along the upper footwall ramp, footwall synclines with compressive deformation of climbing ripple cross-laminated sands are very common, mainly developed as growthfault synclines as mentioned above. all the thrust sheets from gr02 to gr13 can be demonstrated to have been carried piggyback on the gr01 thrust sheet. gr01 thrust sheet the accumulated displacement of gr01 is estimated to about 140 m. this includes an interpreted displacement, c. 40 m, of the thrust-sheet tip. the thrust-sheet tip was, at a post-thrust stage, eroded away by the truncation of the glaciotectonic unconformity; calculation of the displacement of the tip follows the principle illustrated in fig. 11. the total displacement also includes the c. 100 m displacement along the exposed footwall ramp of ss06 and its consequent continuation down to the décollement zone in the 30 m flat level (stratigraphic level from the l/r-unconformity). from the base of the cliff and down into the subsurface, thrusting took place along the gr01 hangingwall flat. the internal tectonic structure of the lønstrup klint formation in gr01 is very similar to the thrust structures described from the ss06 thrust sheet of the stenstue rende section (figs 92, 93) with intraformational duplex structures, imbrication and strong mobilisation along the hanging-wall flat. as can be seen from the cross-section (plate 2b), the l/r-unconformity is traceable down to a level c. 5 m below sea level. this interpretation is supported by field observations, although the base of the clif f is often scree covered, and implies that the hanging-wall flat can be traced down to the décollement surface, and that the thrust sheet has not been elevated up onto, and translated along, intermediate flats in the subsurface. the l/runconformity rests in its initial stratigraphic position, and the reference level lies below sea level. the accufig. 95. large-scale cross-bedding displaying d-onlap overlain by planar bedding (d-onlap) and climbing ripple cross-laminated sand (c-bedding) of the rubjerg knude formation in the grønne rende section. photograph: july 1999; way-up is to the left. l/r-u, l/runconformity. 114 mulated thickness of the rubjerg knude formation is c. 15 m, and the sand beds were deposited with an onlap onto the northerly dipping l/r-unconformity (r-onlap). gr02 thrust sheet the gr02 thrust sheet is the southernmost imbricate in the imbricate fan of the grønne rende section. it consists of a 10 m thick section of the lønstrup klint formation overlain by an about 30 m thick section of the rubjerg knude formation. the main part of the thrust dips at 55°n, whereas the upper part is somewhat steeper. the lower 5–8 m thick unit of the rubjerg knude formation is characterised by large-scale trough crossbedding, and r-onlap can be recognised. f-dipping bedding, grading up into d-onlap in the uppermost part of the clif f section overlies this lower unit; this indicates ramp–flat–foreland dipping relationships during ramp and flat propagation. above this, a middle sand unit with f-bedding was deposited, and finally the upper unit shows r-onlap, which was subsequently folded in a footwall syncline below the gr03 thrust fault. gr03 thrust sheet in the gr03 thrust sheet, the lower part comprising the lønstrup klint formation is a wedge-shaped structure aligned along a vertical thrust fault (gr03 hanging-wall ramp). the thickness varies from c. 15 m in the lower part to c. 3 m in the upper part. above the l/r-unconformity, the rubjerg knude formation consists of a more than 30 m thick succession, which indicates that the piggyback basin of gr03 (as well as gr02) was a long-lived depocentre. the geometry of the gr03 hanging-wall ramp implies that it can be traced down to the 15 m (or 20 m) level. during ramping, a second stage of erosion affected the l/r-unconformity, which thinned out the lønstrup klint formation (fig. 96). the r-onlap in the lower half of the rubjerg knude formation probably reflects the ramping on the footwall ramp of gr02, and the middle part of the rubjerg knude formation was deposited during the propagation of the hanging-wall ramp over the footwall flat of gr02. the uppermost 4 m of the gr03 thrust sheet is very disturbed, probably due to push from the gr04 upper hanging-wall ramp. gr04 thrust sheet in the gr04 thrust sheet, the lønstrup klint formation is relatively thick, c. 20 m in the lower part of the cliff section and about 7 m in the top part. the gr04 hanging-wall ramp dips 70–80°n, and in the middle part of the cliff section a small hanging-wall ramp about 5 m high is preserved. in front of this ramp, the sand deposited at the top of the gr03 piggyback basin was pushed forward during thrust faulting along the upper footwall flat, as mentioned above. fig. 96. the composite development of the l/r-unconformity (l/r-u) resulted in the reduction of the thickness of the gr03 thrust sheet to a very thin horizon interlayered with thick piles of sand referred to the rubjerg knude formation in the grønne rende section. the upper hanging-wall ramp (gr03hwr) displays marked relief due to erosion. photograph: october 2000. 115 the maximum thickness of the rubjerg knude formation in gr04 is similar to the thickness in gr02 and gr03, but the piggyback basin is wedge-shaped due to the footwall ramp produced by the thrusting of gr05. the large-scale trough cross-bedded lower part of the formation tends to show r-onlap towards the upper part of the l/r-unconformity. gr05 thrust sheet gr04 and gr05 initially formed one coherent thrust sheet, with gr05 being carried piggyback during thrust propagation of gr04 before they were separated by the satellite thrusting of gr05 up along the footwall ramp of gr04. the hanging-wall ramp drops from the 5 to the 10 m flat level along a relatively steep ramp, which lifted gr05 free of gr04. in the lower part of the cliff section the l/r unconformity is situated about 5 m above sea level, indicating ramping to an intermediate flat in the subsurface. the gr05 thrust fault is overturned to the north, indicating that the thrust plane dips at 75°s (fig. 97). the uppermost 10–15 m of sand beds in the gr05 piggyback basin are horizontally orientated. the sand beds show large-scale trough cross-bedding and sedimentation is inferred to have taken place between the gr04 and gr05 thrust tips when these were exposed above the sediment/water interface during the latest stage of dynamic development. gr06 thrust sheet the lønstrup klint formation of gr06 is generally a relatively thick unit (10–15 m) although locally in this section deep erosion is evident at the l/r-unconformity. this localised deep erosion at the unconformity was probably due to erosion of a hanging-wall anticline during propagation over the footwall ramp hinge. this suggestion is supported by the presence of r-onlap in the lower to middle part of the cliff. the cfig. 97. the thrust-fault structures related to the gr04 and gr05 thrust sheets. the ‘overturned’ orientation of the gr05 hangingwall ramp (gr05hwr) is regarded as the result of repeated footwall ramping of the gr04 thrust sheet (gr04fwr = footwall ramp of the gr04 thrust sheet), and subsequent translation of the imbricate fan along the lower décollement surface. photograph: july 1999. 116 bedding observed in the middle part of the rubjerg knude formation may well reflect post-ramp deposition; subsequently, sedimentation briefly took place during displacement on the upper footwall flat. the thickness of the rubjerg knude formation is 15 m where the unconformity is deeply incised, decreasing to only 10 m laterally. this indicates that the basin was partly closed by the gr07 thrust propagating along its hanging-wall flat (back of gr06) in an early phase of development of the section. the uppermost beds show a high-angle r-onlap to the unconformity in gr06, demonstrating that the sand was deposited during the final phase of upthrusting, and just before the last c. 30° tilting of the thrust sheet into its present upright position. gr07 thrust sheet the gr07 thrust sheet consists of a 10 m thick unit of the upper part of the lønstrup klint formation with medium-bedded light grey fine-grained sand interbedded with thin mud layers. the l/r-unconformity is parallel with the hanging-wall flat in the main part of the exposed thrust sheet giving the impression that the sheet is of uniform thickness. the topmost part of the sheet is wedge-shaped where the upper hangingwall ramp is preserved, and the irregular structures of the tip can be interpreted as an upper hanging-wall anticline. the rubjerg knude formation in gr07 is about 15 m thick and can be divided into three 5 m thick units, which show the typical characteristics of sedimentation in the formation. in the uppermost part of the piggyback basin, high-angle r-onlap, similar to the bedding in the gr04–gr06 thrust sheets, indicates late syntectonic deposition between the thrust-sheet tips. gr08 thrust sheet the lower part of gr08, comprising the lønstrup klint formation, forms a wedge-shaped structure, with a thickness of only 3 m in the top of the clif f section and about 10 m at the base. the basal part is characterised by mobilised mud bounded at the thrust sole by a vertical thrust fault. at the top, a hanging-wall anticline and small diapir deformed the tip. the rubjerg knude formation of the thrust sheet comprises three units. the lowermost unit, up to 10 m thick, shows r-onlap in the lower part which is also the lower part of the cliff section. upwards, along the steeply dipping l/r-unconformity, the dip of the ronlap increases. there is an angular discordance between these beds and the beds occurring above. these beds show planar parallel bedding (f-bedding). in the lower part of this f-bedded unit, clasts of mud occur with sizes from cobbles to boulders (1 m size). these boulder-sized mud-blocks are interpreted as fragments of the gr09 thrust tip that were deposited by gravity slumping in the piggyback basin. the middle unit of the formation is c. 8 m thick and shows large-scale cross-bedding with sets up to 3 m thick, and the unit has an r-onlap relationship to the unit below. the upper unit is c. 10 m thick and the beds show mainly planar bedding with some trough cross-bedding towards the top. the lower and middle units may be interpreted to represent deposition during two phases of ramp– flat propagation. gr09 thrust sheet the lønstrup klint formation of the gr09 thrust sheet forms a uniform c. 5 m thick unit with a vertical orientation. the rubjerg knude formation is about 25 m thick and can be divided into a lower and an upper part. the lower part displays variable large-scale crossbedding and the unit has a f-bedding relationship, whereas the upper part forms one large r-onlap succession. the lower part may be interpreted as having been deposited while the gr08 hanging-wall ramp propagated over a footwall flat, whereas the upper part was deposited when the gr09 thrust sheet was displaced up along a 45° dipping ramp during a relatively late phase of deformation. a number of minor horizontal extensional faults are interpreted as foreland-dipping normal faults related to a hanging-wall anticline formed over a hanging-wall ramp during an early or intermediate phase of thrusting. gr10 thrust sheet the lønstrup klint formation of gr10 is very thin, only about 3–4 m thick, in the exposed part of the cliff section. the rubjerg knude formation is about 25 m thick, and the lower part shows a poorly exposed f-bedding relationship. the upper part displays marked r-onlap with a 45° dipping angular discordance to the l/r-unconformity. 117 gr11 thrust sheet the gr11 thrust sheet is irregularly orientated, but is mainly vertical in the upper frontal part. in gr11, the lønstrup klint formation has a uniform thickness of 10 m, consisting of medium-bedded light coloured sand interbedded with thin mud layers situated above the steeply dipping hanging-wall flat. the rubjerg knude formation is generally not well exposed due to sand scree, but has a thickness of about 15 m. gr12 thrust sheet a large part of the gr12 thrust sheet is covered by sand scree, and detailed data from this part of the grønne rende section are limited. the lønstrup klint formation forms a c. 5 m thick unit of sand dominated by thick-bedded turbidites, as is typical of the upper part of the formation. the rubjerg knude formation is more than 20 m thick and displays the typical depositional features of the formation. however, it should be noted that the l/r-unconformity in gr12, as is the case in gr11, has been lifted up to an elevation of 15–20 m a.s.l.; this indicates that these sheets propagated over an upper flat, probably composed of two stacked duplex segments of gr06 and gr08 in the subsurface. gr13 thrust sheet the gr13 thrust sheet is also mainly covered by sand scree at the base of cliff. however, exposures in the upper part of the cliff reveal a rather complex structure. the lønstrup klint formation of the thrust sheet is very thin, and in places only the unconformity is observed; it can thus be difficult to recognise where the stratigraphic unconformity is preserved and where it has been completely replaced by the thrust fault, which is now vertically orientated. furthermore, the rubjerg knude formation has been subjected to superimposed folding. the thickness of the piggyback basin deposits in gr13 is more than 30 m and four units of the rubjerg knude formation are differentiated. unit 1 is about 5–6 m thick, and appears in the frontal and upper part of the piggyback basin; it is characterised by large-scale cross-bedding as well as planar parallel stratification. towards the trailing end of the thrust sheet, the r-onlap in unit 1 grades up into unit 2, which is folded into a set of overturned folds, originally with a horizontal axial plane, but now re-orientated into an upright position; the overturned folds deform the bedding in unit 1. unit 3 is characterised by steeply dipping large-scale foresets that were deposited over the recumbent folds. finally unit 4, about 7 m thick, is mainly planar-bedded and was pushed in front of a hanging-wall ramp of the frontal thrust sheet in the rubjerg fyr section. interpretation of structural development two types of interpretations are considered prior to the further description of the structural development: (1) in order to estimate the displacement, a geometric construction has been made for each thrust-sheet tip subsequently eroded away at the glaciotectonic unconformity (fig. 11), and (2) the position of the l/runconformity below the screes has been constructed from successive approximations (plate 1). interpretation of the extent of the thrust tips was carried out as a triangular construction with a best-fit of the intersection of the unconformity and the thrust plane (see fig. 11). where these surfaces (lines in the 2-d constructions in fig. 11) were obscured, the construction was guided by the assumption that the acute angle is close to 18°, which from experience is a general initial thrust ramp angle. according to this geometrical reconstruction, the average displacement of each thrust sheet is about 70 m. the general impression is that the thrusting of the twelve imbricate sheets in the grønne rende fan-structure was broadly contemporaneous. if the experience from the ulstrup section near the foreland is taken into consideration, one would expect a long coherent thrust sheet initially displaced along the upper flat level at 10 m stratigraphic depth (from the l/r-unconformity). from this detachment level, the imbricate thrust-fault fan propagated with fairly equal spacing, and the thrusting progressed during sedimentation of the rubjerg knude formation. however, there are two positions where the piggyback basin has a thinner depositional fill, namely the basin on gr01 and on gr06/gr07. thus, the first longer displacement along an upper footwall flat took place on the back of gr06 and gr01, subsequently leaving a trailing-end duplex segment below the 10 m detachment level at the rear of gr06 as well as at the end of gr01. when about half of the displacement along the gr06 had taken place, the level of detachment started to root down to the 20 m flat level, such that the trailing duplex seg118 ment of gr06 was free to move along intermediate ramps. thus, if about 20 m of the displacement on each of the gr07–gr13 thrusts is accumulated, it amounts to a total displacement of about 140 m for the gr06 trailing segment (gr06u). it is therefore argued that the gr06u segment started to ramp in the middle phase of thrust propagation, which is reflected in a foreland-dipping tilt in the gr13 basin that created the recumbent folding. the sequence in ramping follows three angular modes: initial ramping along an angle of about 18°, intermediate ramping along an angle of 30°, and final ramping at close to 45°. however, ramping will never be initiated at an angle of 90°, so the problem is how to explain the vertically orientated thrust faults. the first c. 20° ramp is given by the upper ramping, and progressive ramping will result in a 36–45° tilt. this would result in progressive steepening of the tilting from south to north, however, which is clearly not the case. it is suggested, therefore, that the gr01 thrust jumped down to the décollement level at the 30 m stratigraphic depth, and that this caused the reorientation during the final displacement of gr01. thus, when the imbricate fan with ramp angles up to 40° was carried along with the lower gr01 thrustsheet segment (gr01u) towards the 45° inclined frontal ramp, all the thrust sheets were subsequently tilted due to a common megascopic shear. the combination of this large-scale shear tilt and the accumulated ramp steepening is very well illustrated in the gr05 thrust sheet, which was separated from gr04 along a satellite thrust fault. the hanging-wall ramp of gr04 was fixed with a c. 75° steep dip. consequently the hanging-wall ramp of gr05, which was carried piggyback on gr04, ended up in an overturned position (dip of about 75°s). in balancing the grønne rende section, one problem remains to be solved, namely the fate of the trailing-end segment of gr06. however, this is a minor problem compared to the space problem created by the gr01 thrusting along the 30 m décollement level. the trailing-end thrust-sheet segment of gr01u (a duplex sheet created between the 30 m and the 20 m flat level) is considered below under the structural and dynamic analysis of the rubjerg knude fyr and stortorn sections. rubjerg knude fyr section this section is situated below the rubjerg knude fyr, and makes up the highest part of the cliff (see cover illustration). the cliff section below the lighthouse comprises the thickest thrust sheets, which are ramped up to the highest footwall flat level. this means that thrust sheets with a décollement level at a depth of about 35 m are ramped up and thrust along a flat at the 10 m deep level. normal fault structures similar to the brede rende normal fault also appear in the rubjerg knude fyr section. the tip of the prominent thrust sheet in the central part of the section was dropped down into the piggyback basin in front of the thrust fault, and subsequently buried by sediments of the rubjerg knude formation. structural elements related to the hanging-wall ramp are well illustrated in this section. tectonic architecture the rubjerg knude fyr section comprises six relatively thick thrust sheets annotated rf01–rf06. the leading-edge thrust and hanging-wall ramp-and-flat of rf01 terminate all the imbricates in the grønne rende section. to the north, the section is bounded by the footwall ramp of rf06 that coincides with the hanging-wall ramp of the frontal thrust sheet in the storntorn section. the thickness of the thrust sheets is bounded by the 20 m flat level in the southern part, and increases in the northern part down to the 30 m flat. all the thrust sheets in the section are thrust over the lower duplex segment of gr06 and gr01. translation along an intermediate flat is indicated by the dominant position of the l/r-unconformity at 20 m a.s.l. the rf01 and rf02 thrust sheets make up an easily recognisable thrust sheet pair that can also be located on the geological cross-section constructed by jessen (1918; fig. 98). the similarity of the remaining part of the rubjerg knude fyr section to jessen’s cross-section is not so obvious, probably due to the intervening 80 years of cliff erosion and the present poor exposure of the section due to extensive sand scree. the dips of the footwall ramps are about 60°. the angle between the l/r unconformity and the thrusts is about 30°, which means that the thrusts have been rotated about 30° during ramp propagation in the subsurface. this is illustrated by the angular relationship between the l/r unconformity in rf01 and the footwall ramp of rf02. the thrust-fault displacement of rf01 and rf02 ranges from 50–65 m. 119 at the tip of rf01, a 25 m long upper hanging-wall flat is preserved. below this, a 5–7 m thick sheet comprising the top of the gr13 piggyback basin occurs. this upper footwall flat segment of gr13 was displaced an unknown distance (25–50 m) forwards in front of the upper hanging-wall ramp of rf01. the rubjerg knude formation in rf02 is poorly exposed due to sand scree at the base of the cliff, but it is interpreted to be 20 m thick. from the l/r-unconformity up to the overlying thrust fault, the thickness of the piggyback basin is about 40 m. however, it is inferred that a thrust fault situated in the middle part of the basin is responsible for repetition of the rubjerg knude formation, and that a normal fault similar to the brnf, is located in the upper levels of the rf02 sheet. the normal fault is asymptotic, fading out towards the l/r-unconformity in the rf02 thrust sheet. the rf03 thrust sheet is a small sheet with a truncation structure. the nose of this thrust sheet was obviously exposed to normal faulting at an early stage of development (fig. 99). after fault displacement, the rf03 tip was eroded away and the top of the thrust sheet erosionally truncated to form an unconformity, which cut off the sheet at a very steep angle (> 70°). the normal fault at the tip of rf03 is the first of two normal faults displaced down on to the piggyback basin of rf02. in the second phase of normal faulting, the c. 45 m long tip of the rf04 thrust sheet was displaced c. 35 m down a normal fault plane (fig. 100), which had an angle of 70–90° relative to the thrust fault and the l/r-unconformity of rf04. the normal faulting may have been initiated by differential translation of hanging-wall ramps along intermediate flats related to one or more subsurface duplex segments. in a late phase of thrust faulting, the trailing part of rf04 was thrust up over the piggyback basin of rf04/rf03/rf02, which brought the nearly 45° dipping lower hanging-wall ramp of rf04 into contact with the footwall ramp of the displaced tip of rf04. the final phase of fault-bend folding tilted this ramp into a 70°s dipping position. the rf05 and rf06 thrust sheets were thrust up fig. 98. the thrust sheet pair exposed below the rubjerg knude fyr. to the right, the rf01 thrust sheet is thrust faulted along its hanging-wall ramp (rf01hwr) along the footwall ramp in the trailing end of the grønne rende section. rf01 forms the footwall block for the propagation of the hanging-wall ramp of rf02 (rf02hwr). note the elevated position of the l/r-unconformity indicating that the thrust sheets were thrusted up on duplex segments in the subsurface. photograph: july 1999. geological survey of denmark and greenland bulletin 41, 2018, 83-86 83 arctic sea ice has a significant impact on the global radiation budget, oceanic and atmospheric circulation and the stability of the greenland ice sheet (vaughan et al. 2013). prior to the era of aircraft and satellite, information on sea-ice extent relied on observations from ships and people living at the coast. this information is a valuable contribution to better understand the history of sea ice. however, the information exists in a range of formats, e.g., sea-ice extent before the late 1800s is typically reported in the literature as an annual index from a single geographical point or as hand-drawn maps. this makes it difficult to assess and compare data across time and space. the combination of digitised historical maps and single-point data makes the information more accessible and provides a record that can help understand the dynamics and processes of the climate and its interactions with the cryosphere (chapman & walsh 1993). in this study, maps of sea-ice extent by koch (1945) were digitised. we use these maps in combination with sea-ice charts from the danish meteorological institute (dmi) and koch’s sea-ice index from 1820 to 1939, to map estimated sea-ice extent between iceland and greenland going back to 1821. this information has not been included in even the most recent databases of arctic sea ice (walsh et al. 2015, 2017). furthermore, we extract time series of sea-ice extent at a number of locations and investigate the relationship between them. our observation area is along eastern greenland, between the southern tip of greenland at 59°46́ n northwards to 77°21́ n. digitising lauge koch’s maps koch’s (1945) monograph comprises over 400 small maps that show monthly sea-ice extent from 1877 to 1939. we digitised the maps for the months in those years that contained sea ice. the data used by koch were compiled by thoroddsen (1884) from annals that describe the sea-ice extent off iceland. the final result is 134 digital maps showing sea-ice extent from 1877 to 1939. these maps are available for download at http://dx.doi.org.pangaea.de/10.1594/pangaea.887453 reconstructing the extent of sea ice between greenland and iceland from koch’s index koch’s monograph contains a sea-ice index that categorises the width of the ice belt off iceland, based on observations from 1820 to 1939. the koch index has three main categories from a to c and subcategories from 1 to 3 for each letter (table 1). a represents a narrow ice belt year, b is a broad ice belt year and c depicts a very broad ice belt year. the number next to the letter indicates how far the sea ice spread off iceland. a subcategory value of 1 means the ice did not reach iceland, 2 means it had spread along the northern coast and 3 indicates it had spread down as far as the southern coast of iceland. koch described an a year as the least severe sea-ice observationally constrained reconstruction of 19th to mid-20th century sea-ice extent off eastern greenland danielle a.m. hallé, nanna b. karlsson, anne munck solgaard and camilla s. andresen ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! !! ! ! ! ! ! !! ! ! ! ! ! ! ! ! ! ! ! ! ! greenland iceland 0 500 1,000250 km %100-80 %79-60 ! ! %59-40 n o r t h a t l a n t i c o c e a n n probability of sea ice present based on the years 1893-1956 minimum and maximum extent for a1 years fig. 1. example of the final map, available for downloading, showing the probability of sea-ice extent based on historical maps from 1893 to 1956. © 2018 geus. geological survey of denmark and greenland bulletin 41, 83–86. open access: www.geus.dk/bulletin http://dx.doi.org.pangaea.de/10.1594/pangea.887453 http://dx.doi.org.pangaea.de/10.1594/pangea.887453 http://www.geus.dk/bulletin 8484 year, whereas c is the most severe sea-ice year, with b being of intermediate severity. for example, the year 1895 is classified as b3 because it had a wider area of sea ice that touched on northern iceland. table 1 gives a summary of the index for each year. the duration of sea ice is not taken into consideration in these indexes. the dataset we produced comes from three different sources of historical maps. the first is the digitised icelandic maps, the second source is the dmi sea-ice charts that we digitised for 1893 to 1900 for the months between march and september each year. we have only digitised sea-ice extent along the east coast to the southern tip of greenland from the dmi charts. these charts were based on ice conditions from variable sources such as land-based observations, scientific explorations and ship logs. the third source is shapefiles from dmi that are available online. these files contain digitised sea-ice data from 1901 to 1956 (underhill et al. 2014). we subsequently used these data to create the digitised maps. the dmi shapefiles were downloaded from the nsidc at: https://nsidc.org/data/ g10007/versions/1. the original maps used to create the shapefiles from 1893 to 1956 can be found on the nsidc website: https://nsidc.org/data/g02203/versions/1#. due to world war ii, there is a gap in the data from 1939 to 1946. digitisation of the maps was done using arcmap and exported as completed map files in jpeg format. next, and for ease of data manipulation, the images were processed to convert the values into binary arrays, so that sea ice has a value of 1 and no sea ice has a value of 0. this was done in order to better visualise and plot the data. for each category (a1, a2...c) of the index, where there are at least 10 years present (table 1) we created a map showing the probability that sea ice had a certain extent. the maps contain three lines of sea ice that depict 80–100%, 60–79% and 40–59% probability of sea-ice extent (fig. 1). the percentages are calculations of the average sea-ice extent. we used the average sea-ice extent for each year belonging to each category to calculate the frequency of the extent. for example, the dataset shows that in 80–100% of those years within a certain category the sea ice extended to the thick dashed line, 60–79% of the time the extent reached the solid black line and only 40–59% of the time did the sea ice extend to the thin dashed line. the result of this procedure is six maps (a1, a2, a3, b1, b2 and c) from 1893 to 1956. in some cases, we can go even further back in time; three maps showing iceland only (b2, b3 and c) have been constructed for the years 1877 to 1939. figure 2 shows an example of a c year. there are no probability maps for a years for the iceland-only maps, because of a lack of data. n iceland n o r t h a t l a n t i c o c e a n 100-80% 79-60% 59-40% 0 150 30075 km probability of sea ice extent based on the years 1877-1939 maximum and minimum extents for c years fig. 2. example of probability map of sea ice off iceland 1877 to 1939. index year 1884 1823 1853 1924 1830 1902 1862 1913 1836 1916 1821 1861 1914 1822 1868 1911 1922 1832 1867 1928 1831 1903 1876 1838 1917 1824 1870 1915 1828 1869 1918 1926 1839 1899 1930 1835 1923 1877 1850 1921 1825 1875 1919 1837 1878 1927 1841 1904 1933 1854 1925 1879 1880 1932 1827 1883 1938 1840 1881 1931 1843 1905 1937 1864 1929 1889 1885 1829 1886 1855 1882 1934 1844 1908 1871 1889 1890 1834 1895 1856 1887 1935 1845 1909 1872 1893 1897 1857 1896 1858 1888 1936 1846 1910 1873 1894 1898 1859 1906 1865 1891 1939 1848 1920 1901 1900 1912 1860 1907 1866 1892 total table 1. summary of how koch’s index was classified for each year from 1821 to 1939 a1 a2 a2 a2 a3 a3 b1 b1 b2 b2 b3 b3 b3 c c c 9 23 14 10 13 22 20 https://nsidc.org/data/g10007/versions/1. https://nsidc.org/data/g10007/versions/1. https://nsidc.org/data/g02203/versions/1 85 for further analyses of the sea-ice data, 12 points were selected (fig. 3) to pinpoint areas of interest in order to investigate if there is a relationship between the points. a conditional probability matrix (table 2) was created to analyse the relationship between the 12 selected points using the digitised datasets described above. we calculated the probability that if there is sea ice at a given point then there is sea ice at another point. for example, the value 0.177 in the last column indicates that if point 12 has sea ice then there is a 0.177 probability that point 11 also has sea ice. a strong probability is defined as a value of 0.850 to 1.000, a medium probability is defined as a value of 0.490 to 0.849 and a weak probability is less than 0.490. results in this section, we discuss the observed relationship between the east greenland sea ice and the sea ice off iceland using nine probability maps and the conditional probability matrix (table 2). koch describes an a1 year as a year when the ice between greenland and iceland would have been narrow and iceland free of ice throughout that year. during a1 years the probability maps show that sea ice would definitely have been present along the coast of greenland, down to the southern tip with fluctuations out from the coast but the ice would not have reached iceland. in a2 years, sea ice is approaching the north coast of iceland, but iceland itself would have been almost completely free of ice. it is further estimated that 80–100% of the time, sea ice would most likely have been at the northern part of the coast, and for only 60–79% of the time to be present at the southern part of iceland. in a3 years the sea-ice belt would have been wider at times extending into the northern fjords of iceland. in b years, the sea ice would have formed a broad belt north of iceland. koch describes in his report that the ice edge reached near 67°n and 15°w, which corroborates the digitised data. in b1 years, the probability maps reflects a wider range of sea-ice extent along the coast of greenland. for example, at scoresby sund (point 7), the maps show that the fjord is not always covered by ice during those years. in b2 years, the probability map shows that the probability of 80 to 100% covers a larger area along greenland’s east coast, and has an extent of sea ice that reached further along the greenland 70°n 60°n 70°n 60°n 10°e 50°w 40°w 30°w 20°w 10°w 30°e60°w iceland faroe islands svalbard 1 5 6 10 9 12 7 8 114 3 2 fig. 3. map of the location of the twelve points used for the probability matrix. 1 2 3 4 5 6 7 8 9 10 11 12 1 1.000 0.916 0.896 0.942 0.789 0.625 0.889 0.898 0.880 0.969 0.828 0.169 2 0.924 1.000 0.948 0.942 0.842 0.500 0.937 0.933 0.930 0.984 0.841 0.171 3 0.896 0.940 1.000 0.919 0.789 0.625 0.913 0.918 0.907 0.977 0.832 0.168 4 0.976 0.976 0.960 1.000 0.842 0.500 0.980 0.988 0.981 0.977 0.845 0.165 5 0.060 0.064 0.060 0.062 1.000 0.750 0.063 0.063 0.062 0.094 0.082 0.316 6 0.020 0.016 0.020 0.015 0.316 1.000 0.016 0.016 0.016 0.031 0.026 0.375 7 0.900 0.940 0.924 0.950 0.842 0.500 1.000 0.961 0.946 0.961 0.823 0.163 8 0.920 0.948 0.940 0.969 0.842 0.500 0.972 1.000 0.965 0.984 0.845 0.169 9 0.912 0.956 0.940 0.973 0.842 0.500 0.968 0.976 1.000 0.984 0.828 0.167 10 0.498 0.502 0.502 0.481 0.632 0.500 0.488 0.494 0.488 1.000 0.509 0.305 11 0.771 0.777 0.775 0.754 1.000 0.750 0.758 0.769 0.744 0.922 1.000 0.177 12 0.977 1.000 0.977 1.000 0.140 0.070 0.953 1.000 1.000 0.907 0.953 1.000 table 2. a conditional probability matrix known point with sea ice the left hand column indicates the point with sea ice (in relation to fig. 2), while the rows show the probability of sea ice present relative to that point pr ob ab ilit y po in t h as se a ice 8686 icelandic northern coast. unlike that observed in b1 years, for b2 years scoresby sund was often filled with sea ice. finally, during c years, sea ice was present along the entire east greenland coast and for 40–59% of the years the ice would extend down the eastern coast of iceland. of the nine maps, the remaining three show sea-ice data off iceland from 1877 to 1939. these maps provide more details on sea-ice extent during a more severe ice period in the 1880s (kelly et al. 1987, based largely on koch’s index of icelandic sea-ice occurrence). in b2 years, there was a 40 to 59% probability of sea ice reaching all along the northern coast of iceland, but persisting mainly between iceland and greenland. the b3 years have a wider area between greenland and iceland and cover a further range along the northern icelandic coast. kelly et al. (1987) discuss three of the b3 years (1914, 1915 and 1938) as being years with extensive sea ice around iceland. during c years, the sea ice extended to the south coast of iceland. seven of these years (1881, 1882, 1887, 1888, 1892, 1911 and 1918) were severe sea-ice years for iceland (kelly et al. 1987). a statistical analysis was carried out to see if there is any correlation between sea ice being present at one point (table 2) to other points on the map. the strongest probabilities are the points along eastern greenland, points 1–4 and 7–9. this implies that if there is sea ice at one point along the coast then it is very likely that the rest of the eastern coast of greenland would have been filled with ice. point number 12 is located farther from the coast but has a perfect correlation with points 2, 4, 8 and 9 along the coast. the points farther from the coast 10–12 have a range of weak to strong probability with the other points. point 12 also shows a relatively higher probability value of 0.316 and 0.375 with points 5 and 6 than the other points closer to greenland, which all have weak values of less than 0.094. lauge koch’s index extends further back in time than the maps of the spatial extent. when using the probability maps, it is now possible to see how extensive the sea ice was likely to have been in the period from 1821 to 1893. the final application of the maps is to use them in combination with table 1 to be able to estimate sea-ice extent back to 1821. conclusion 134 maps showing sea-ice extent from 1877 to 1939 are available as jpeg files and shapefiles. detailed probability maps have been created from observed and estimated sea-ice conditions from observed data that date back to the early 1800s. these probability maps can be downloaded at https://dx.doi. org/10.1594/pangaea.887452 where the maps are available as jpeg files. the datasets can be used in further studies to help understand the characteristics of sea ice during the 19th and 20th centuries. previous studies have used koch’s sea-ice index to infer atmospheric circulation patterns (kelly et al. 1987). our new dataset can provide important input to studies of climatic variability in the (sub)arctic including the link between ice and atmospheric circulation modes such as the arctic and north atlantic oscillations (ao/nao) and thus helps predicting how the climate may respond to future sea-ice loss (screen et al. 2017). references chapman, w.l. & walsh, j.e. 1993: recent variations of sea ice and air temperature in high latitudes. bulletin of the american meteorological society 74, 33–47. kelly, p.m., goodess, c.m. & cherry, b.s.g. 1987: the interpretation of the icelandic sea ice record. journal of geophysical research: oceans 92(c10), 10835–10843. koch, l. 1945: the east greenland ice. meddelelser om grønland 130(3), 373 pp. screen, j.a. 2017: the missing northern european winter cooling response to arctic sea ice loss. nature communications 8, 14603. thoroddsen, t. 1884: den grönlandska drifisen vid island. ymer 1884, 145–160. underhill, v., fetterer, f. & petersen, c. 2014: arctic sea ice concentration and extent from danish meteorological institute sea ice charts, 1901–1956, version 1. boulder: nsidc (national snow and ice data center), https://dx.doi.org/10.7265/n5mp517m vaughan, d.g. et al. 2013: observations: cryosphere. in: stocker, t.f. et al. (eds): climate change 2013: the physical science basis. contribution of working group i to the fifth assessment report of the intergovernmental panel on climate change, 317–382. cambridge: cambridge university press. walsh, j.e., chapman, w.l. & fetterer, f. 2015, updated 2016: gridded monthly sea ice extent and concentration, 1850 onwards, version 1.1, boulder: national snow and ice data center. digital media, http:// dx.doi.org/10.7265/n5833pz5 walsh, j.e., fetterer, f., stewart, j.s. & chapman, w.l. 2017: a database for depicting arctic sea ice variations back to 1850. geographical review 107, 89–107, http://dx.doi.org/10.1111/j.1931-0846.2016.12195.x authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: dhalle23@gmail.com. https://doi.pangaea.de/10.1594/pangaea.887452 https://doi.org/10.7265/n5mp517m http://dx.doi.org/10.7265/n5833pz5 http://dx.doi.org/10.7265/n5833pz5 http://dx.doi.org/10.1111/j.1931-0846.2016.12195.x mailto:dhalle23@gmail.com geological survey of denmark and greenland bulletin 3, 1-23 1 geological survey of denmark and greenland bulletin 3 · 2004 late quaternary environmental changes recorded in the danish marine molluscan faunas kaj strand petersen geological survey of denmark and greenland ministry of the environment geus bulletin no 3.pmd 28-06-2004, 08:451 2 geological survey of denmark and greenland bulletin 3 keywords bottom-communities, climate changes, danish, environment, interglacial–glacial cycle, late quaternary, marine, mollusc faunas. cover donax vittatus on the sandy shores of northern france. kaj strand peteresen danmarks og grønlands geologiske undersøgelse øster voldgade 10, dk-1350 copenhagen k, denmark e-mail: ksp@geus.dk scientific editor of this volume: svend stouge editorial secretaries: esben w. glendal and birgit eriksen referees: svend funder and gotfred høpner petersen, denmark illustrations: gurli e. hansen bengaard and henrik klinge pedersen digital photographic work: benny m. schark and jakob lautrup graphic production: knud gr@phic consult, odense, denmark printers: schultz grafisk, albertslund, denmark manuscript submitted: 9 january 1998 final version approved: 11 december 2003 printed: 15 july 2004 isbn 87-7871-122-1 geological survey of denmark and greenland bulletin the series geological survey of denmark and greenland bulletin replaces geology of denmark survey bulletin and geology of greenland survey bulletin. citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 3, 268 pp. available from geological survey of denmark and greenland øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk or geografforlaget aps rugårdsvej 55, dk-5000 odense c, denmark phone: +45 63 44 16 83, fax: +45 63 44 16 97, e-mail: go@geografforlaget.dk © danmarks og grønlands geologiske undersøgelse (geus), 2004 geus bulletin no 3.pmd 28-06-2004, 08:452 3 contents abstract . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7 introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9 danish sites with marine sediments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9 the late pleistocene . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10 the holocene . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 12 the recent fauna of shell-bearing molluscs compared to the subfossil fauna . . . . . . . . . . . 12 molluscan finds within the seven regions during the holocene . . . . . . . . . . . . . . . . . . . . . 15 the bælt sea area . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 15 the baltic area . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16 the kattegat area with fjords . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16 the limfjord area . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17 the north sea . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19 the vendsyssel area . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 20 the skagen well area . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 22 the danish late quaternary marine molluscs . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 24 class polyplacophora . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25 order neoloricata . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25 class gastropoda . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25 subclass prosobranchia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25 order archaeogastropoda . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25 order mesogastropoda . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 28 order heterogastropoda . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 39 order neogastropoda . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 44 subclass heterobranchia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 50 order heterostropha . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 50 subclass opisthobranchia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 55 order bullomorpha . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 55 order anaspidea . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 58 order thecosomata . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 59 order gymnosomata . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 60 subclass pulmonata . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 60 order basommatophora . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 60 class scaphopoda . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 60 order siphonodentalioida . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 60 order dentalioida . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 61 class bivalvia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 62 subclass palaeotaxodonta . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 62 order nuculoida . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 62 subclass pteriomorphia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 66 order arcoida . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 66 order mytiloida . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 67 order pteroida . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 69 subclass heterodonta . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 73 order veneroida . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 73 order myoida . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 92 subclass anomalodesmata . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 95 order pholadomyoida . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 95 geus bulletin no 3.pmd 28-06-2004, 08:453 4 the skagen well . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 98 the skagen well – perspectives . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 98 the pre-late quaternary deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 99 the late pleistocene . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 99 eemian deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 99 the early/middle weichselian, marine and glacigene deposits . . . . . . . . . . . . . . . . . . 100 the late weichselian marine and glacigene deposits . . . . . . . . . . . . . . . . . . . . . . . . . 101 the holocene . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 102 the preboreal–boreal 10 000 – 8000 14c years b.p. . . . . . . . . . . . . . . . . . . . . . . . . . . 103 the atlantic 8000–5000 14c years b.p. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 104 the subboreal 5000–2500 14c years b.p. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 105 the subatlantic 2500– 14c years b.p. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 105 the older subatlantic . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 106 the younger subatlantic . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 107 conclusive remarks on the skagen well . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 112 the environmental changes through time in the seven sectors based on the molluscan records . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 112 eemian species sorted after climatic affinities . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 112 the bælt sea . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 112 the baltic . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 116 the kattegat . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 118 the north sea . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 120 vendsyssel . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 122 skagen . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 125 early/middle weichselian species sorted after climatic affinities . . . . . . . . . . . . . . . . . . . . 126 the kattegat . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 126 vendsyssel . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 127 skagen . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 129 late weichselian species sorted after climatic affinities . . . . . . . . . . . . . . . . . . . . . . . . . 130 vendsyssel . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 130 skagen . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 132 holocene species sorted after climatic affinities . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 133 the bælt sea . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 133 the baltic . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 135 the kattegat . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 138 the limfjord . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 140 the north sea . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 143 vendsyssel . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 146 skagen . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 149 the environmental changes within the seven regions through the late quaternary evaluated by the molluscan communities met with in the seven stages . . . . . . . . . . . . 151 eemian stage 130 000 – 115 000 b.p. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 152 the bælt sea, region 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 152 the baltic, region 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 154 the kattegat, region 3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 154 the north sea, region 5 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 155 the vendsyssel region, region 6 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 156 the skagen region, region 7 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 156 early/middle weichselian stage 115 000 – 25 000 b.p. . . . . . . . . . . . . . . . . . . . . . . . . . 157 the kattegat, region 3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 157 the vendsyssel region, region 6 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 157 geus bulletin no 3.pmd 28-06-2004, 08:454 5 the skagen region, region 7 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 157 late weichselian stage 25 000 – 10 000 b.p. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 158 the vendsyssel region, region 6 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 158 the skagen region, region 7 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 159 the preboreal–boreal stage 10 000 – 8000 14c years b.p. . . . . . . . . . . . . . . . . . . . . . . . 159 the north sea, region 5 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 159 the skagen region, region 7 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 159 the atlantic stage 8000–5000 14c years b.p. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 160 the bælt sea, region 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 160 the baltic, region 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 160 the kattegat, region 3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 160 the limfjord, region 4 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 161 the north sea, region 5 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 161 the vendsyssel region, region 6 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 161 the skagen region, region 7 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 163 the subboreal stage 5000–2500 14c years b.p. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 163 the bælt sea, region 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 163 the baltic, region 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 163 the kattegat, region 3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 164 the limfjord, region 4 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 164 the north sea, region 5 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 166 the vendsyssel region, region 6 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 166 the skagen region, region 7 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 167 the subatlantic stage 2500– 14c years b.p. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 167 the bælt sea, region 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 168 the baltic, region 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 168 the kattegat, region 3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 168 the limfjord, region 4 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 169 the north sea, region 5 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 169 the vendsyssel region, region 6 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 170 the skagen region, region 7 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 171 concluding remarks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 172 acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 172 references . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 173 list of synonyms . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 180 index of species . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 188 appendix 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 197 recent species . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 197 subfossil species . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 201 recent species sorted after climatic affinities . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 204 subfossil species sorted after climatic affinities . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 208 appendix 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 213 appendix 3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 218 appendix 4 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 258 appendix 5 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 260 appendix 6 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 264 geus bulletin no 3.pmd 28-06-2004, 08:455 6 geus bulletin no 3.pmd 28-06-2004, 08:456 7 abstract petersen, k.s. 2004: late quaternary environmental changes recorded in the danish marine molluscan faunas. geological survey of denmark and greenland bulletin 3, 268 pp. late quaternary, marine deposits in denmark have yielded 247 subfossil species of molluscs. the sites are presented, and comparisons are made between the subfossil mollusc assemblages and the 278 shell-bearing mollusc species presently living in the danish seas. 184 species are common to the two groups. the 63 species no longer occurring around denmark are used as indicators of changing environmental conditions, including temperature, salinity and depth, throughout the last 130 000 years. seven modern faunal regional units are defined and considered: the bælt, the baltic, the kattegat, the limfjord, the north sea and the vendsyssel regions, and the skagen area based on the skagen iii well dgu file no. 1.287. the late quaternary, marine, shell-bearing molluscs, comprising 341 subfossil and recent species, are characterised from the point of view of climatic (i.e. arctic, subarctic, boreal and lusitanian) affinities and animal–sediment relationships. on this background the faunal and environmental evolution recorded in the 217 m long skagen well core is analysed and described. the mollusc assemblages in the skagen sequence indicate a deeper-water facies during the eemian, the weichselian and the older holocene in contrast to what hitherto was known in other parts of the danish area during the late quaternary. for the skagen well the chronozones preboreal/boreal, atlantic, subboreal and subatlantic can be identified by 14c dating. the environmental changes within the seven regions through the late quaternary are evaluated by depicting the molluscan communities encountered in the seven late quaternary stages together with remarks on studies of the neighbouring areas. by following the marine communities through the late quaternary in the light of the classical bottom communities sensu c.g.j. petersen, it is demonstrated how facies have changed both through time and space within the danish marine realm. the wellestablished, more temperate eemian marine fauna was closely associated with shallow-water environments. the inferred climatic changes reflect an interglacial–glacial cycle. however, the climatically induced changes during the holocene in the marine environment were small and overshadowed by the facies changes. out of the 341 species recorded in this study, 140 occur in the eemian, 36 in the early/middle weichselian and 41 in the late weichselian. the holocene fauna is represented by 183 species of shell-bearing molluscs, of which the first recorded occurrence of 148 species has been radiocarbon-dated. author’s address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ksp@geus.dk geus bulletin no 3.pmd 28-06-2004, 08:457 8 50 km korsør holbæk høng blåvands huk mandø hølade frederikshavn korupsø agger tange agger bovbjerg ertebølle tastum sø gedser darss stavtrup hals hollerup ejby bro limhamn aarhus stensigmose voderup klint ristinge strandegaard dyrehave hobro aalbæk jydske rev saltholm læsø anholt skagen amager djursland vendsyssel dybvad skærumhede ærø jerup vester holmen samsø højen vust kovad bro vognsbøl fredericia esbjerg varde forballum grærup farup tønder yder bjerrum røjle klint holmstrup røsnæs ulfborg bulbjerg løkken strandby bindslev bornholm møn rügen vejle fjord mariager fjord tybrind vig sidinge fjord skive fjord nissum fjord isefjord roskilde fjord limfjorden fakse bugt skagerrak lille bæ lt store bæ lt ø resund københavn jylland sjælland fyn north sea baltic sea bælt sea kattegat 57° 55° 8° 10° 12° 14° fig. 1. location map with late quaternary marine localities and names of areas on land and of danish waters. geus bulletin no 3.pmd 28-06-2004, 08:458 9 introduction in the middle of 19th century, denmark had its first ‘geology of denmark’ published by g. forchhammer, in 1835. however, as forchhammer expressed it in 1851 when making some notes on the work by the malacologist o.a.l. mørch (1828–1878) at the mineralogical museum of copenhagen. it has hitherto been enough for the geognost to establish formations using the characteristic fossils, but in the future we have to give a closer description from a zoological point of view (petersen 1997, p. 5). considering only the younger deposits, the efforts of the zoologist in geological works are highly significant and became important already in the 19th century. c.g.j. petersen (1860–1928) is an outstanding example of such an influence with his work on the extent of shell-bearing molluscs in the danish seas inside the skagen (petersen 1888, 1893). here he points to the faunal conditions also in the pleistocene and holocene marine deposits compared to the recent distribution. in the description accompanying the geological map sheets of vendsyssel (jessen 1899), jessen gives full credit to c.g.j. petersen and a. jensen (1866– 1953) for their studies on the molluscan species recorded from that part of the country. later both petersen and jensen contributed further to our knowledge of the marine molluscan fauna. petersen formed the concept of the bottom communities (petersen & jensen 1911; petersen 1913, 1914, 1915, 1918) that has been the tool for further work, not only within the danish waters but all over the world with the so-called parallel bottom communities (thorson 1957). though the concept of parallel molluscan communities in the sense of thorson (1957) has been considerably modified in the last 30 years (erwin 1983), there remains a recognition that particular molluscan assemblages are associated with various types of habitat. in 1899 the zoologist v. nordmann (1872–1962) was engaged by the geological survey of denmark to study the molluscs from the quaternary deposits. part of this work was already reflected in the next geological map sheet covering the southern part of vendsyssel (jessen 1905). here nordmann has identified the molluscs and given the faunal remarks on the holocene marine fauna in the north-eastern part of the limfjord (fig. 1). in his work, the zoological considerations are given, elucidating the holocene palaeoenvironments. however, from the beginning of the century nordmann touched upon many other aspects within the late quaternary marine environments which form the most important base for the present study covering marine deposits from the eemian, the weichselian and the holocene. in the following chapter the presentation of some observed sites with marine sediments will be given as an introduction to an answer to the question raised by petersen (1910, p. 29): “what i have often missed in the geological studies is a thorough or detailed comparison between the fossil faunas and the molluscan faunas now living before our eyes”. the aim of this work is to characterise the changing environments in the danish waters through time as seen in the macrofaunas and bottom communities mainly based on molluscs. danish sites with marine sediments initially, the findings and descriptions of the danish marine localities shown in fig. 1 were part of the university studies pioneered by g. forchhammer. however, since the start of the geological survey of denmark in 1888, much of the information has come from the systematic mapping of denmark, and the results have been published in the descriptions to the geological map sheets of denmark (fig. 2). as seen from the plan for the geological mapping of denmark (e.g. sørensen & nielsen 1978) it was decided to do the mapping first in the northern parts of jylland and sjælland and to present a record of the marine deposits from the areas mapped. today, up to 80 per cent of the country has been mapped and descriptions for many map sheets have been published. the main information on the holocene marine molgeus bulletin no 3.pmd 28-06-2004, 08:459 10 luscs is available in these publications and is used in the present description supplemented by specific molluscan studies within the areas. consequently, the frame will be the transition area between the north sea and the baltic and the descrip57° 55° 8° 10° 12° 14° 50 km 1. bælt sea 2. baltic sea 3. kattegat vendsyssel 7. skagen n o r t h s e a 4. limfjorden 5. 6. region 1 – 7 fig. 2. the frame for the seven regions follows mainly the pattern of the old geological map sheets (sørensen & nielsen 1978, fig. 1) and partly the regions used in jensen & knudsen (1995, fig. 1). 1: the bælt sea covering the southern part of the bælts. 2: the baltic covering the southern part of øresund and east of darss–gedser. 3: the kattegat region covering the northern part of the bælts and øresund. 4: the western limfjord – except the north sea coastal region. 5: the north sea with coastal regions and skagerrak. 6: vendsyssel including former marine areas. 7: skagen, mainly the skagen well dgu file no. 1.287. tion mainly based on the geological map sheets found in the following regions shown in fig. 2: (1) the bælt sea; (2) the baltic; (3) the kattegat with bordering fjords; (4) the limfjord; (5) the north sea; (6) vendsyssel and the skagen well iii, dgu file no. 1, 287. the late pleistocene in 1841 forchhammer found the cyprina clay to the southern part of denmark, naming the unit after the dominating bivalve (forchhammer 1842). first, however, forchhammer referred the thick shell molluscs to glossus humanus rather than to arctica islandica. consequently, he placed the deposits in the ‘brunkulsformation’, viz. the tertiary. when finally realising that the common species was arctica islandica, he transferred the deposits to the so-called ‘rullestensformation’, viz. the quaternary. along with the investigations of the cyprina clay through the years since 1841, the actual stratigraphical position was very much under debate, and it was not until 1928 when nordmann wrote his la position stratigraphique des dépôts d’eem that the cyprina clay attained its final position: “appartenant à la dernière période interglaciaire” (nordmann 1928, p. 65). later the name ‘eemian’ became the designation for the whole interglacial, according to gripp (1964, pp. 215–216). johnstrup (1882a) gave the first detailed description of the cyprina clay in denmark and slesvig. also in the northern part of denmark, late pleistocene deposits were studied by johnstrup (1882b), but with references to the earlier works by forchhammer (1822), bredsdorff (1824), faber (1828) and pingel (1828). in 1908 nordmann made his doctoral thesis on the molluscan fauna from the cyprina clay and other central european deposits, forming a part of the publication by madsen et al. (1908). the sequence of interglacial–glacial marine deposgeus bulletin no 3.pmd 28-06-2004, 08:4510 11 its is described from the well at skærumhede (jessen et al. 1910). here the full late pleistocene record is found, although the stratigraphic position was not clear at that time. later investigations, also with studies of the molluscan fauna, were published in 1974 and a late pleistocene age proposed (bahnson et al. 1974). the difference between the boreo-lusitanian community in the boring and the typical eemian community as found in southern denmark was interpreted as difference in facies (bahnson et al. 1974) (see nilsson 1983). in the study of the marine late pleistocene deposits in southern denmark (ødum 1933) based on the record of molluscan species as determined by v. nordmann the finds point to two different deposits in time. one is regarded as eemian and the other as the so-called skærumhede fauna. however, later investigations at strandegaards dyrehave in southern sjælland (petersen & konradi 1974) and at holmstrup in central sjælland (fig. 1; petersen & buch 1974), revealed that the molluscan species found at strandegaards dyrehave, one of the localities of ødum (1933) and regarded as representing the skærumhede fauna, could be eemian but reflecting another facies than the typical eemian on the islands south of fyn. the holmstrup fauna is to be correlated with the arctic marine weichselian in northern jylland which is the upper part of the portlandia arctica zone sensu nordmann (madsen et al. 1908) or the macoma calcarea zone sensu petersen (bahnson et al. 1974, fig. 7), see fig. 3 for stratigraphical position. the aminostratigraphic investigations of the danish late pleistocene deposits as published by miller & mangerud (1985) sustain only to some extent the abovementioned correlations: “none of the sites regarded here as eemian (strandegaards dyrehave) gave ratios as high as in holsteinian deposits or as low as in middle weichselian deposits” (miller & mangerud 1985, p. 261). in the case of the holmstrup weichselian site, only three out of eleven individuals of macoma calcarea gave weichselian ratios (miller & mangerud 1985, p. 264). the marine molluscan fauna of the late weichselian has been studied intensively only from the vendh ol oc en e la te w ei ch se lia n pl ei st oc en e m id dl e w ei ch se lia n ea rl y w ei ch . eemian l a t e q u a t e r n a r y subatlantic subboreal atlantic boreal preboreal younger dryas allerød older dryas bølling historical age iron age bronze age neolithic mesolithic palaeolithic young baltic swedish old baltic norwegian 1000 2000 3000 4000 5000 6000 7000 8000 9000 10000 11000 12000 13000 14000 15000 16000 17000 18000 19000 20000 21000 22000 1050 2050 2900 3700 4400 5500 6100 7200 8100 9100 10000 10200 11100 12100 12800 13400 14200 15100 16100 17000 17600 18500 25000 75000 115000 130000 period epoch age culture stagechron ice stream chronology calendar years bp 14c years bp fig. 3. stratigraphic framework for the late quaternary deposits from pedersen & petersen (1997). geus bulletin no 3.pmd 28-06-2004, 08:4511 12 syssel area recorded in the publications by jessen (1899, 1936). nearly 30 molluscan samples from these late weichselian – younger yoldia sea deposits have been dated (krog & tauber 1974). theevaluationof themolluscancommunities invendsyssel reveals the changing late weichselian sea level (petersen 1984), and the highest marine shoreline, around 60 m a.s.l., in northern denmark can be shown to develop between 14 000 and 13 000 b.p. (14c years). the holocene forchhammer participated in the work of the so-called ‘lejrekomité’, an interdisciplinary committee studying human remains along the shore. this commission gave the first – and now famous – description of the ‘køkkenmødding’ (kitchen midden), a mound consisting of shells of edible molluscs and other refuse, marking the site of a prehistoric human habitation (hanks 1971). ‘køkkenmødding’ is one of the few danish international terms (forchhammer et al. 1851). the work of the ‘lejrekomité’ was concentrated on the marine molluscs in order to establish out whether the shell deposits were naturally based – oyster banks – or whether they were formed as waste deposits produced by men living at coastal sites. the other members of the commission were j. worsaae and j. steenstrup, representing archaeology and zoology respectively. consistently, the study of the molluscan elements was based mainly on steenstrup’s work. however, while working in the commission, forchhammer continued his studies on the sea levels (forchhammer 1838, 1840). this was essential for the discussion of whether the molluscs found belonged to raised marine deposits or were gathered by man. forchhammer’s study led to the concept of raised marine deposits north of a line from nissum fjord to south of korsør in the storebælt area (fig. 1). this line still carries the name of forchhammer and divides the country into two parts, with the raised marine areas to the north-east, and to the south-west the area where the land has been sinking. together with the study of the holocene molluscan fauna by johnstrup (1882b), such observations on shorelines were also collected. it became one of the points specially mentioned in the instructions for the autographic geologists when the systematic geological mapping of denmark was started in 1888 by the geological survey of denmark (sørensen & nielsen 1978). the recent fauna of shell-bearing molluscs compared to the subfossil fauna the record of recent danish shell-bearing molluscs has been taken from the annotated check list of recent marine molluscs of danish waters (jensen & knudsen 1995). in appendix 1 the species are presented taxonomically following jensen & knudsen (1995). late immigrants from the last centuries – transferred by man -– have been omitted from the list, because the aim of the present study is to present the development in the subfossil late quaternary molluscan fauna also in appendix 1 compared to the natural fauna of today. according to fredén (1986), subfossil means that the weight of the object when found does not exceed its original weight, which is obviously the case for younger deposits seen geologically as shells from the late quaternary. in all, 278 recent species of shell-bearing molluscs are recorded from the danish waters: the class polyplacophora is represented by seven geus bulletin no 3.pmd 28-06-2004, 08:4512 13 species forming 2.5% of the total number of known species. the class gastropoda is represented by 151 species forming 54.3% of the total number of known species. the class scaphopoda is represented by three species forming 1.1% of the total number of known species. the class bivalvia is represented by 117 species forming 42.1% of the total number of known species. the list of known finds of subfossil species amounts to 247 species. with regard to the classes, it appears that polyplacophora is now represented by only one species, which formed 0.4% of the total subfossil molluscan species. within the class gastropoda 125 species occur, forming 50.6% of the total number of subfossil species, a figure which is nearly 5% lower than that for recent gastropods. the class scaphopoda is represented by five fossil species which form 2.0% of the subfossil shell-bearing species which is a little higher than the ratio for the recent fauna. the class bivalvia is represented by 116 species forming 47.0% of the total, which is a little more than 7% above the recent ratio. the low number of subfossil polyplacophora can be explained by the fact that the shells from those species are nearly always broken, and this excludes identification to species level, so to say, following the statement made by knudsen (1970, p. 1): “isolated and worn plates were neglected altogether”. among the gastropods, the subclasses and orders, except the order heterostropha within the subclass heterobranchia, have a lower representation of subfossil finds than of recent ones. the heterostropha, which has a 2.5% higher representation among the subfossil finds than among the recent ones, is a group fig. 4. regional division of the european seas from feyling-hanssen (1955). geus bulletin no 3.pmd 28-06-2004, 08:4513 14 of mostly tiny specimens which might be more looked for in the geological samples than in the recent bottom samples often used in the more practical work of evaluation benthos introduced by c.g.j.petersen. however, many of these small species should be considered with the utmost care, with respect to the difficulty of identifying them to species level within subfossil material. the reason why the class scaphopoda has a twice as great a representation within the subfossil material cannot be given, although it is tempting to regard the different palaeoenvironment back in the late quaternary as the explanation of the higher frequency. the greater variety of palaeoenvironment and different climate back in time is clearly the reason why the bivalvia within all subclasses has a higher percentage than in the recent fauna. however, as an overview, the total subfossil species could be compared to the recent ones arranged also after their climatic affinities, as will be thoroughly discussed in one of the following chapters. with respect to distribution of molluscan species within the north atlantic – west european realm, four zones may be distinguished, viz.: the arctic = a, the subarctic = s, the boreal = b and the lusitanian = l (figs 4, 5). it appears from the comparison between subfossil species and recent species sorted after climatic affinity (appendix 1) that the subfossil species have their dominance in the extreme groups, i.e. arctic = a; arctic/ subarctic = as; arctic, subarctic and boreal = asb and subarctic/boreal, while the species with a wide tolerance – arctic, subarctic, boreal and lusitanian = asbl – have a higher representation within the recent fauna. also the middle group, which is represented by faunal element from the subarctic, boreal and lusitanian, the boreal and lusitanian (which is the most numerous group with 140 subfossil species) has a higher representation in the recent fauna. but the group of purely lusitanian species has a clearly better representation among the subfossil species, as seen by the percentage figure 6.2% compared to 0.7% for the purely lusitanian faunal elements among the subfossil and recent faunas respectively. these observations reveal that the late quaternary fauna covers a period of 130 000 years with changing climatic conditions both with colder and warmer periods than at present. so considering the totals of subfossil and recent species one has to discuss the difference not only quantitatively but qualitatively; because only 184 species are shared between the late quaternary and the recent finds, while 63 species have to be considered as particular ones occurring within the late quaternary during the eemian, the weichselian or the holocene, in one, two or in all three groups but not the recent one. within the bivalvia, the highest amount of subfossil species (31) found only in the late quaternary occur. such species are also the species which must be focused on in the evaluation of the changing environment through time. fig. 5. regional division of the european seas according to símonarson et al. (1998). geus bulletin no 3.pmd 28-06-2004, 08:4514 15 molluscan finds within the seven regions during the holocene the molluscan finds within each region (see fig. 2) from the holocene, as appearing mainly from the descriptions accompanying the geological map sheets of denmark, are presented. the bælt sea area from the bælt sea area the information on the occurrences of molluscs has been taken from the following map sheets: madsen (1902) and jessen (1907 (contributions by v.nordmann), 1935, 1945); v.milthers (1940) and k. milthers (1959). nordmann (1906) has a record of molluscs found in skælskør nor (sw sjælland) and petersen records from the areas south of fyn, storebælt and lillebælt (1985c, 1989). subfossil holocene species in the bælt sea area class gastropoda subclass prosobranchia order neotaenioglossa littorina littorea (linnaeus 1758) littorina obtusata (linnaeus 1758) littorina saxatilis (olivi 1792) littorina tenebrosa (montagu 1803) lacuna pallidula (da costa 1778) lacuna vincta (montagu 1803) hydrobia ulvae (pennant 1777) hydrobia ventrosa (montagu 1803) onoba semicostata (montagu 1803) rissoa albella lovén 1846 rissoa inconspicua alder 1844 rissoa membranacea (j. adams 1800) bittium reticulatum (da costa 1778) order heterogastropoda triphora adversa (montagu 1803) order neogastropoda hinia reticulata (linnaeus 1758) subclass heterobranchia order heterostropha omalogyra atomus (phillippi 1841) odostomia conoidea winckworth 1932 subclass opisthobranchia order anaspidea retusa obtusa (montagu 1803) retusa truncatula (bruguière 1792) akera bullata müller 1776 subclass pulmonata order basommatophora lymnaea peregra (müller 1774) class bivalvia subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 modiolula phaseolina (philippi 1844) modiolus modiolus (linnaeus 1758) musculus discors (linnaeus 1767) order pterioida ostrea edulis linnaeus 1758 subclass heterodonta order veneroida mysella bidentata (montagu 1803) tridonta borealis schumacher 1817 parvicardium exiguum (gmelin 1791) parvicardium ovale (sowerby 1840) parvicardium scabrum (philippi 1844) cerastoderma edule (linnaeus 1758) spisula subtruncata (da costa 1778) angulus tenuis (da costa 1778) macoma balthica (linnaeus 1758) scrobicularia plana (da costa 1778) abra alba (wood 1802) arctica islandica (linnaeus 1767) paphia aurea (gmelin 1791) tapes decussatus (linnaeus 1758) venerupis pullastra (montagu 1803) order myoida mya arenaria linnaeus 1758 mya truncata linnaeus 1758 corbula gibba (olivi 1792) hiatella arctica (linnaeus 1758) barnea candida (linnaeus 1758) zirfaea crispata (linnaeus 1758) total for the holocene bælt sea: 47 (19.0%) geus bulletin no 3.pmd 28-06-2004, 08:4515 16 the baltic area the baltic area is here restricted to the area east of darss and south of øresund at saltholm, which must be considered the baltic sensu stricto when regarding the present distribution of the marine fauna and also taking into consideration the subfossil holocene molluscan fauna, as will be demonstrated by a following comparison with the other areas. the main map sheet published is by v. milthers from 1908 with contributions by v. nordmann on the holocene molluscan fauna. the subfossil holocene fauna has also been studied later in the western part by petersen (1994b). in the description accompanying the map sheet bornholm (grönwall & milthers 1916) there is no record of a mollusc fauna. subfossil holocene species in the baltic area class gastropoda subclass prosobranchia order neotaenioglossa littorina littorea (linnaeus 1758) littorina tenebrosa (montagu 1803) hydrobia ulvae (pennant 1777) hydrobia ventrosa (montagu 1803) rissoa albella lovén 1846 rissoa inconspicua alder 1844 rissoa membranacea (j. adams 1800) bittium reticulatum (da costa 1778) aporrhais pespelicani (linnaeus 1758) order neogastropoda hinia reticulata (linnaeus 1758) subclass opisthobranchia order anaspidea retusa truncatula (bruguière 1792) subclass pulmonata order basommatophora lymnaea peregra (müller 1774) class bivalvia subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 subclass heterodonta order veneroida parvicardium exiguum (gmelin 1791) cerastoderma edule (linnaeus 1758) cerastoderma glaucum (poiret 1789) macoma balthica (linnaeus 1758) scrobicularia plana (da costa 1778) order myoida corbula gibba (olivi 1792) total for the holocene baltic: 19 (7.7%) the kattegat area with fjords the kattegat area sensu lato includes the fjords, i.e. the northern part of the lillebælt area, storebælt and øresund. therefore the following map sheets are taken within this area: 1. the north-eastern part of sjælland described by rørdam (1893), who published a detailed description of the holocene marine deposits from northeast sjælland already in 1891 and continued with the description of the map sheet københavn and roskilde (1899) where the southernmost parts of the roskilde fjord and the øresund are described in great detail for the marine holocene part. 2. furthermore, rørdam & v. milthers published the description for the geological map sheet of nw sjælland in 1900 and nordmann on the molluscs in sidinge fjord (westerby 1933). 3. from the north-western part of fyn and the island of samsø by madsen (1897, 1900) and together with ussing for the north-eastern part of fyn (ussing & madsen 1897). 4. the map sheet of fredericia (nordmann 1958) covers the northern part of the lillebælt and vejle fjord on the eastern coast of jylland. 5. from the islands of læsø and anholt in the kattegat the description was given by jessen (1897) and nordmann (1903a). 6. from hobro by nordmann (jessen 1927). 7. from mariager fjord by nordmann (ødum 1929). 8. the peninsula of djursland has been mapped during the last decades, and the description of the holocene marine molluscan fauna is by petersen (pedersen & petersen 1997) and petersen (1993). subfossil holocene species in the kattegat class gastropoda subclass prosobranchia order archaeogastropoda theodoxus fluviatilis (linnaeus 1758) order neotaenioglossa littorina littorea (linnaeus 1758) littorina obtusata (linnaeus 1758) littorina saxatilis (olivi 1792) littorina tenebrosa (montagu 1803) lacuna vincta (montagu 1803) hydrobia ulvae (pennant 1777) hydrobia ventrosa (montagu 1803) skeneopsis planorbis (fabricius 1780) onoba semicostata (montagu 1803) onoba vitrea (montagu 1803) geus bulletin no 3.pmd 28-06-2004, 08:4516 17 rissoa albella lovén 1846 rissoa inconspicua alder 1844 rissoa membranacea (j. adams 1800) bittium reticulatum (da costa 1778) order heterogastropoda triphora adversa (montagu 1803) order neogastropoda buccinum undatum linnaeus 1758 hinia reticulata (linnaeus 1758) subclass heterobranchia order heterostropha omalogyra atomus (phillippi 1841) brachystomia eulimoides hanley 1844 chrysallida spiralis (montagu 1803) subclass opisthobranchia order anaspidea retusa obtusa (montagu 1803) retusa truncatula (bruguière 1792) akera bullata müller 1776 class bivalvia subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 musculus discors (linnaeus 1767) order pterioida heteranomia squamula (linnaeus 1758) ostrea edulis linnaeus 1758 subclass heterodonta order veneroida mysella bidentata (montagu 1803) acanthocardia echinata (linnaeus 1758) parvicardium exiguum (gmelin 1791) parvicardium scabrum (philippi 1844) cerastoderma edule (linnaeus 1758) cerastoderma glaucum (poiret 1798) macoma balthica (linnaeus 1758) scrobicularia plana (da costa 1778) abra alba (wood 1802) arctica islandica (linnaeus 1767) paphia aurea (gmelin 1791) tapes decussatus (linnaeus 1758) venerupis pullastra (montagu 1803) order myoida mya arenaria linnaeus 1758 mya truncata linnaeus 1758 corbula gibba (olivi 1792) hiatella rugosa (linnaeus 1758) total for the holocene kattegat: 45 (18.2%) the limfjord area from the limfjord area (western part), excluding the part which falls within vendsyssel, only one description for a map sheet has been published (gry 1979). however, the molluscs are recorded in publications by petersen (1976, 1981, 1985a, 1986a) and in rasmussen & petersen (1980). furthermore, v.nordmann collected holocene marine shells from the western limfjord in 1902–1903 which were further examined by erna nordmann and leifur símonarson in the sixties as mentioned in petersen (1976, p. 78). it must be emphasised that c.g.j. petersen in 1888 discussed the subfossil fauna also from the limfjord, which was earlier the topic of collin (1884). subfossil holocene species in the limfjord class gastropoda subclass prosobranchia order archaeogastropoda patella vulgata linnaeus 1758 helcion pellucidum (linnaeus 1758) iothia fulva (müller 1776) acmaea tessulata (müller 1776) acmaea virginea (müller 1776) margarites helicinus (phipps 1774) gibbula cineraria (linnaeus 1758) gibbula tumida (montagu 1803) skenea serpuloides (montagu 1808) skenea basistriata (jeffreys 1877) order neotaenioglossa littorina littorea (linnaeus 1758) littorina obtusata (linnaeus 1758) littorina saxatilis (olivi 1792) littorina tenebrosa (montagu 1803) lacuna pallidula (da costa 1778) lacuna parva (montagu 1803) lacuna vincta (montagu 1803) hydrobia ulvae (pennant 1777) hydrobia ventrosa (montagu 1803) skeneopsis planorbis (fabricius 1780) alvania lactea (michaud 1830) alvania punctura (montagu 1803) cingula semistriata (montagu 1808) onoba semicostata (montagu 1803) onoba proxima (forbes & hanley 1850) onoba vitrea (montagu 1803) rissoa albella lovén 1846 rissoa inconspicua alder 1844 rissoa membranacea (j. adams 1800) rissoa parva (da costa 1779) rissoa violacea desmarest 1814 geus bulletin no 3.pmd 28-06-2004, 08:4517 18 caecum glabrum (montagu 1803) bittium reticulatum (da costa 1778) turritella communis risso 1826 aporrhais pespelicani (linnaeus 1758) lunatia alderi (forbes 1838) lunatia catena (da costa 1778) order heterogastropoda triphora adversa (montagu 1803) cerithiopsis barleei (jeffreys 1867) cerithiopsis tubercularis (montagu 1803) epitonium clathrus (linnaeus 1758) epitonium turtonis (turton 1819) aclis minor (brown 1827) vitreolina philippii (rayneval & ponzi 1854) order neogastropoda nucella lapillus (linnaeus 1758) buccinum undatum linnaeus 1758 hinia incrassata (ström 1768) hinia pygmaea (lamarck 1822) hinia reticulata (linnaeus 1758) oenopota turricola (montagu 1803) raphitoma purpurea (montagu 1803) raphitoma linearis (montagu 1803) subclass heterobranchia order heterostropha omalogyra atomus (phillippi 1841) brachystomia eulimoides hanley 1844 odostomia scalaris macgillivray 1843 chrysallida decussata (montagu 1803) chrysallida eximia (jeffreys 1849) chrysallida indistincta (montagu 1808) chrysallida obtusa (brown 1827) chrysallida spiralis (montagu 1803) ebala nitidissima (montagu 1803) eulimella laevis (brown 1827) eulimella scillae (scacchi 1835) ondina divisa (j. adams 1797) ondina diaphana (jeffreys 1848) odostomia acuta jeffreys 1848 odostomia conoidea winckworth 1932 odostomia turrita hanley 1844 odostomia albella lovén 1846 odostomia plicata (montagu 1803) turbonilla crenata (brown 1827) turbonilla delicata (monterosato 1874) turbonilla lactea (linnaeus 1758) subclass opisthobranchia order bullomorpha acteon tornatilis (linnaeus 1758) cylichna cylindracea (pennant 1777) cylichna alba (brown 1827) philine aperta (linnaeus 1767) philine punctata (adams 1800) order anaspidea diaphana minuta brown 1827 retusa obtusa (montagu 1803) retusa truncatula (bruguière 1792) retusa umbilicata (montagu 1803) akera bullata müller 1776 class bivalvia subclass palaeotaxodonta order nuculoida nucula nitidosa winckworth 1930 nucula nucleus (linnaeus 1767) nuculoma tenuis (montagu 1808) subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 modiolula phaseolina (philippi 1844) modiolus adriaticus (lamarck 1819) modiolus modiolus (linnaeus 1758) musculus discors (linnaeus 1767) modiolaria tumida (hanley 1843) order pterioida aequipecten opercularis (linnaeus 1758) chlamys varia (linnaeus 1758) delectopecten vitreus (gmelin 1791) palliolum striatum (müller 1776) palliolum tigerinum (müller 1776) pododesmus patelliformis (linnaeus 1761) anomia ephippium linnaeus 1758 heteranomia squamula (linnaeus 1758) ostrea edulis linnaeus 1758 subclass heterodonta order veneroida lucinoma borealis (linnaeus 1758) thyasira flexuosa (montagu 1803) mysella bidentata (montagu 1803) tellimya ferruginosa (montagu 1803) turtonia minuta (fabricius 1780) lepton nitidum (turton 1822) tridonta borealis schumacher 1817 acanthocardia echinata (linnaeus 1758) parvicardium exiguum (gmelin 1791) parvicardium ovale (sowerby 1840) parvicardium scabrum (philippi 1844) cerastoderma edule (linnaeus 1758) cerastoderma glaucum (poiret 1789) mactra stultorum (linnaeus 1758) lutraria lutraria (linnaeus 1758) spisula elliptica (brown 1827) spisula solida (linnaeus 1758) geus bulletin no 3.pmd 28-06-2004, 08:4518 19 spisula subtruncata (da costa 1778) ensis ensis (linnaeus 1758) phaxas pellucidus (pennant 1777) angulus tenuis (da costa 1778) fabulina fabula (gmelin 1791) macoma balthica (linnaeus 1758) donax vittatus (da costa 1778) gari fervensis (gmelin 1791) scrobicularia plana (da costa 1778) abra alba (wood 1802) abra nitida (müller 1776) abra prismatica (montagu 1803) arctica islandica (linnaeus 1767) chamelea striatula (da costa 1778) clausinella fasciata (da costa 1778) paphia aurea (gmelin 1791) tapes decussatus (linnaeus 1758) timoclea ovata (pennant 1777) venerupis rhomboides (pennant 1777) venerupis pullastra (montagu 1803) mysia undata (pennant 1777) order myoida mya truncata linnaeus 1758 corbula gibba (olivi 1792) hiatella arctica (linnaeus 1758) saxicavella jeffreysi winckworth 1930 barnea candida (linnaeus 1758) pholas dactylus linnaeus 1758 zirfaea crispata (linnaeus 1758) subclass anomalodesmata order pholadomyoida thracia phaseolina (lamarck 1818) total for the holocene limfjord: 147 (59.5%) the north sea in the north sea region the map sheet blaavands huk (fig. 1) forms the southernmost part of what is covered by the present presentation regarding the holocene deposits, and this area was described by jessen (1925). nordmann (in jessen 1925) contributed with the study of the molluscs. 1. in the work by petersen (1985a) the molluscan fauna in the coastal region – the aggertange – is recorded. 2. the geological map sheet from ulfborg was published by petersen et al. (1992a), and the molluscan fauna treated by petersen, but not yet published, is included. 3. in 1994 the holocene molluscs from the jydske rev were studied and reported in a work for the danish coastal authority (petersen 1994a), and with minor corrections published in petersen (1998). subfossil holocene species in the north sea class gastropoda subclass prosobranchia order archaeogastropoda gibbula cineraria (linnaeus 1758) theodoxus fluviatilis (linnaeus 1758) order neotaenioglossa littorina littorea (linnaeus 1758) littorina obtusata (linnaeus 1758) littorina saxatilis (olivi 1792) lacuna pallidula (da costa 1778) lacuna parva (montagu 1803) lacuna vincta (montagu 1803) hydrobia ulvae (pennant 1777) hydrobia ventrosa (montagu 1803) cingula turgida (jeffreys 1870) onoba vitrea (montagu 1803) rissoa albella lovén 1846 rissoa inconspicua alder 1844 rissoa membranacea (j. adams 1800) rissoa violacea desmarest 1814 caecum glabrum (montagu 1803) bittium reticulatum (da costa 1778) turritella communis risso 1826 aporrhais pespelicani (linnaeus 1758) lunatia alderi (forbes 1838) lunatia catena (da costa 1778) order heterogastropoda triphora adversa (montagu 1803) epitonium clathrus (linnaeus 1758) aclis ascaris (turton 1819) aclis minor (brown 1827) aclis walleri jeffreys 1867 order neogastropoda nucella lapillus (linnaeus 1758) buccinum undatum linnaeus 1758 hinia pygmaea (lamarck 1822) hinia reticulata (linnaeus 1758) subclass heterobranchia order heterostropha brachystomia eulimoides hanley 1844 chrysallida indistincta (montagu 1808) chrysallida spiralis (montagu 1803) eulimella laevis (brown 1827) ondina diaphana (jeffreys 1848) odostomia conoidea winckworth 1932 geus bulletin no 3.pmd 28-06-2004, 08:4519 20 odostomia albella lovén 1846 turbonilla crenata (brown 1827) turbonilla delicata (monterosato 1874) turbonilla lactea (linnaeus 1758) subclass opisthobranchia order bullomorpha acteon tornatilis (linnaeus 1758) cylichna alba (brown 1827) order anaspidea retusa obtusa (montagu 1803) retusa truncatula (bruguière 1792) retusa umbilicata (montagu 1803) class bivalvia subclass palaeotaxodonta order nuculoida nucula nitidosa winckworth 1930 nucula nucleus (linnaeus 1767) nucula sulcata (bronn 1831) nuculoma tenuis (montagu 1808) subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 musculus discors (linnaeus 1767) order pterioida chlamys varia (linnaeus 1758) heteranomia squamula (linnaeus 1758) ostrea edulis linnaeus 1758 subclass heterodonta order veneroida thyasira flexuosa (montagu 1803) mysella bidentata (montagu 1803) tellimya ferruginosa (montagu 1803) lepton nitidum (turton 1822) acanthocardia echinata (linnaeus 1758) parvicardium exiguum (gmelin 1791) parvicardium ovale (sowerby 1840) parvicardium scabrum (philippi 1844) parvicardium minimum (philippi 1836) cerastoderma edule (linnaeus 1758) mactra stultorum (linnaeus 1758) spisula elliptica (brown 1827) spisula solida (linnaeus 1758) spisula subtruncata (da costa 1778) ensis ensis (linnaeus 1758) phaxas pellucidus (pennant 1777) angulus tenuis (da costa 1778) fabulina fabula (gmelin 1791) macoma balthica (linnaeus 1758) donax vittatus (da costa 1778) scrobicularia plana (da costa 1778) abra alba (wood 1802) abra nitida (müller 1776) arctica islandica (linnaeus 1767) chamelea striatula (da costa 1778) clausinella fasciata (da costa 1778) paphia aurea (gmelin 1791) tapes decussatus (linnaeus 1758) timoclea ovata (pennant 1777) venerupis pullastra (montagu 1803) dosinia exoleta (linnaeus 1758) dosinia lincta (montagu 1803) order myoida mya truncata linnaeus 1758 corbula gibba (olivi 1792) hiatella arctica (linnaeus 1758) saxicavella jeffreysi winckworth 1930 barnea candida (linnaeus 1758) zirfaea crispata (linnaeus 1758) subclass anomalodesmata order pholadomyoida cochlodesma praetenue (pulteney 1799) thracia phaseolina (lamarck 1818) total for the holocene north sea: 95 (38.5%) the vendsyssel area the vendsyssel area includes the description accompanying the map sheets over the northern, central and southern parts, all by jessen (1899, 1905), but with a contribution by v. nordmann, who wrote the part on the holocene molluscan fauna in the latter publication. in this description by nordmann he presents the different faunal communities as discovered in the subfossil assemblages. it was nordmann’s intention to continue the work further west into the western limfjord area, but his first investigations were not used in the systematic geological mapping. they were, however, of great importance for the understanding of the development of the holocene molluscan fauna (nordmann 1910, 1918). in 1928 in connection with the international congress in copenhagen a final overview by nordmann of the quaternary marine deposits in denmark was given in the summary of the geology of denmark (madsen et al. 1928). here nordmann points to the dosinia layers first described at the beginning of the century from vendsyssel (nordmann 1904), with a record of a fauna not found in the older tapes beds originally demonstrated by petersen (1888). geus bulletin no 3.pmd 28-06-2004, 08:4520 21 later investigations by lauersen (1937) and petersen (1990,1991a, b, 1992) on the dosinia beds at strandby are also included in the list of holocene marine molluscs from vendsyssel. in just pedersen’s thesis on holocene molluscs from 1976 (unpublished), a new fauna element donax vittatus in the dosinia beds is recorded from frederikshavn. subfossil holocene species in the vendsyssel class gastropoda subclass prosobranchia order archaeogastropoda helcion pellucidum (linnaeus 1758) acmaea virginea (müller 1776) gibbula cineraria (linnaeus 1758) gibbula tumida (montagu 1803) order neotaenioglossa littorina littorea (linnaeus 1758) littorina obtusata (linnaeus 1758) littorina saxatilis (olivi 1792) littorina tenebrosa (montagu 1803) lacuna pallidula (da costa 1778) lacuna parva (montagu 1803) lacuna vincta (montagu 1803) hydrobia ulvae (pennant 1777) skeneopsis planorbis (fabricius 1780) alvania lactea (michaud 1830) alvania cimicoides (forbes 1844) alvania punctura (montagu 1803) cingula semistriata (montagu 1808) onoba semicostata (montagu 1803) onoba vitrea (montagu 1803) rissoa albella lovén 1846 rissoa inconspicua alder 1844 rissoa membranacea (j. adams 1800) rissoa parva (da costa 1779) rissoa violacea desmarest 1814 caecum glabrum (montagu 1803) bittium reticulatum (da costa 1778) turritella communis risso 1826 aporrhais pespelicani (linnaeus 1758) trivia monacha (da costa 1778) lunatia alderi (forbes 1838) lunatia catena (da costa 1778) order heterogastropoda triphora adversa (montagu 1803) epitonium clathrus (linnaeus 1758) epitonium turtonis (turton 1819) vitreolina philippii (rayneval & ponzi 1854) order neogastropoda nucella lapillus (linnaeus 1758) buccinum undatum linnaeus 1758 neptunea antiqua (linnaeus 1758) hinia incrassata (ström 1768) hinia pygmaea (lamarck 1822) hinia reticulata (linnaeus 1758) cytharella coarctata (forbes 1840) oenopota turricola (montagu 1803) raphitoma linearis (montagu 1803) subclass heterobranchia order heterostropha omalogyra atomus (phillippi 1841) brachystomia eulimoides hanley 1844 odostomia scalaris macgillivray 1843 chrysallida indistincta (montagu 1808) chrysallida obtusa (brown 1827) chrysallida spiralis (montagu 1803) ebala nitidissima (montagu 1803) eulimella laevis (brown 1827) odostomia conoidea winckworth 1932 odostomia turrita hanley 1844 odostomia albella lovén 1846 odostomia plicata (montagu 1803) turbonilla lactea (linnaeus 1758) subclass opisthobranchia order bullomorpha acteon tornatilis (linnaeus 1758) cylichna cylindracea (pennant 1777) philine aperta (linnaeus 1767) philine punctata (adams 1800) order anaspidea diaphana minuta brown 1827 retusa obtusa (montagu 1803) retusa truncatula (bruguière 1792) retusa umbilicata (montagu 1803) akera bullata müller 1776 class bivalvia subclass palaeotaxodonta order nuculoida nucula nitidosa winckworth 1930 nucula nucleus (linnaeus 1767) subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 modiolula phaseolina (philippi 1844) modiolus adriaticus (lamarck 1819) modiolus modiolus (linnaeus 1758) musculus discors (linnaeus 1767) modiolaria tumida (hanley 1843) order pterioida aequipecten opercularis (linnaeus 1758) geus bulletin no 3.pmd 28-06-2004, 08:4521 22 chlamys varia (linnaeus 1758) pecten maximus (linnaeus 1758) pododesmus patelliformis (linnaeus 1761) anomia ephippium linnaeus 1758 heteranomia squamula (linnaeus 1758) ostrea edulis linnaeus 1758 subclass heterodonta order veneroida lucinoma borealis (linnaeus 1758) thyasira flexuosa (montagu 1803) mysella bidentata (montagu 1803) tellimya ferruginosa (montagu 1803) turtonia minuta (fabricius 1780) lepton nitidum (turton 1822) kellia suborbicularis (montagu 1803) acanthocardia echinata (linnaeus 1758) parvicardium exiguum (gmelin 1791) parvicardium ovale (sowerby 1840) parvicardium scabrum (philippi 1844) cerastoderma edule (linnaeus 1758) laevicardium crassum (gmelin 1791) mactra stultorum (linnaeus 1758) lutraria lutraria (linnaeus 1758) spisula elliptica (brown 1827) spisula solida (linnaeus 1758) spisula subtruncata (da costa 1778) ensis ensis (linnaeus 1758) phaxas pellucidus (pennant 1777) angulus tenuis (da costa 1778) fabulina fabula (gmelin 1791) macoma balthica (linnaeus 1758) macoma calcarea (gmelin 1791) donax vittatus (da costa 1778) gari depressa (pennant 1777) gari fervensis (gmelin 1791) scrobicularia plana (da costa 1778) abra alba (wood 1802) abra nitida (müller 1776) abra prismatica (montagu 1803) arctica islandica (linnaeus 1767) chamelea striatula (da costa 1778) clausinella fasciata (da costa 1778) paphia aurea (gmelin 1791) tapes decussatus (linnaeus 1758) timoclea ovata (pennant 1777) venerupis rhomboides (pennant 1777) venerupis pullastra (montagu 1803) dosinia exoleta (linnaeus 1758) dosinia lincta (montagu 1803) mysia undata (pennant 1777) order myoida mya arenaria linnaeus 1758 mya truncata linnaeus 1758 corbula gibba (olivi 1792) hiatella arctica (linnaeus 1758) hiatella rugosa (linnaeus 1758) saxicavella jeffreysi winckworth 1930 barnea candida (linnaeus 1758) pholas dactylus linnaeus 1758 zirfaea crispata (linnaeus 1758) subclass anomalodesmata order pholadomyoida thracia phaseolina (lamarck 1818) total for holocene vendsyssel: 133 (53.8%) the skagen well area the hitherto recorded molluscan assemblages from danish deposits of late quaternary age are littoral to sublittoral – mostly – especially from the holocene. the new information from the skagen well containing deeper-water deposits is presented below. subfossil holocene species in the skagen well class gastropoda subclass prosobranchia order neotaenioglossa lacuna pallidula (da costa 1778) hydrobia ulvae (pennant 1777) barleeia unifasciata (montagu 1803) onoba vitrea (montagu 1803) rissoa albella lovén 1846 rissoa violacea desmarest 1814 bittium reticulatum (da costa 1778) turritella communis risso 1826 aporrhais pespelicani (linnaeus 1758) lunatia alderi (forbes 1838) lunatia montagui (forbes 1838) order heterogastropoda epitonium trevelyanum (johnston 1841) aclis minor (brown 1827) polygireulima sinuosa (sacco 1836) vitreolina collensi (sykes 1903) vitreolina philippii (rayneval & ponzi 1854) graphis albida (kanmacher 1798) melanella lubrica (monterosato 1891) melanella alba (da costa 1778) hemiaclis ventrosa (jeffreys ms fricle 1874) order neogastropoda buccinum undatum linnaeus 1758 geus bulletin no 3.pmd 28-06-2004, 08:4522 23 hinia pygmaea (lamarck 1822) hinia reticulata (linnaeus 1758) oenopota turricola (montagu 1803) mangelia brachystoma (philippi 1844) subclass heterobranchia order heterostropha chrysallida decussata (montagu 1803) eulimella scillae (scacchi 1835) odostomia conoidea winckworth 1932 odostomia umbilicaris (malm 1863) turbonilla delicata (monterosato 1874) turbonilla sinuosa (jeffreys 1884) subclass opisthobranchia order bullomorpha cylichna alba (brown 1827) order anaspidea retusa truncatula (bruguière 1792) retusa umbilicata (montagu 1803) class bivalvia subclass palaeotaxodonta order nuculoida nucula nitidosa winckworth 1930 nucula nucleus (linnaeus 1767) nuculana minuta (müller 1776) subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 musculus discors (linnaeus 1767) order pterioida chlamys varia (linnaeus 1758) heteranomia squamula (linnaeus 1758) ostrea edulis linnaeus 1758 subclass heterodonta order veneroida thyasira flexuosa (montagu 1803) mysella bidentata (montagu 1803) tellimya ferruginosa (montagu 1803) turtonia minuta (fabricius 1780) acanthocardia echinata (linnaeus 1758) parvicardium minimum (philippi 1836) mactra stultorum (linnaeus 1758) spisula subtruncata (da costa 1778) phaxas pellucidus (pennant 1777) angulus tenuis (da costa 1778) fabulina fabula (gmelin 1791) tellina pygmaea (lovén 1846) donax vittatus (da costa 1778) gari fervensis (gmelin 1791) abra alba (wood 1802) abra nitida (müller 1776) abra prismatica (montagu 1803) arctica islandica (linnaeus 1767) chamelea striatula (da costa 1778) timoclea ovata (pennant 1777) order myoida mya arenaria linnaeus 1758 corbula gibba (olivi 1792) hiatella arctica (linnaeus 1758) saxicavella jeffreysi winckworth 1930 barnea candida (linnaeus 1758) pholas dactylus linnaeus 1758 subclass anomalodesmata order pholadomyoida lyonsia norvegica (gmelin 1791) cochlodesma praetenue (pulteney 1799) thracia phaseolina (lamarck 1818) total for holocene skagen well: 71 (28.7%) geus bulletin no 3.pmd 28-06-2004, 08:4523 geological survey of denmark and greenland bulletin 26, 2012, 41 efforts to include geological and geodetic observations in the assessment of earthquake activity in denmark søren gregersen and peter h. voss assessment of earthquake hazard is improved if geological and geodetic data are included in addition to seismological data. in earthquake regions like japan and california, palaeoseismology combines data from geology and seismology, and networks of permanent gps (global positioning system) stations situated on bedrock supplement the networks of seismographs. a combination of seismographs and gps stations on ice is also used in current studies of glacial earthquakes in greenland. in denmark only broad-scale geodetic coverage is available (khan et al. 2005) and only a few examples of geological input are found in the literature. however, more thorough geodetic evaluations of deformations are currently made in nordic co-operation projects and we are looking forward to learn about the results. also the number of permanent gps stations in denmark has recently been increased from 3 to 13. the geological input is limited but may hold some potential, and the aim of this article is to evaluate this. we also discuss new investigations of recent geological movements, both in projects on post-glacial uplift (and accompanying horizontal deformation) of scandinavia, and more locally of geological indicators of uplift in selected areas. as seismologists, we are interested in a homogeneous evaluation of geological indicators in all of denmark and its neighbouring areas. in two recent seismological papers, we emphasised the very low earthquake activity in denmark, even in the geologically significant sorgenfrei–tornquist zone (gregersen & voss 2009, 2010). the recorded earthquake activity over fig. 1. map of north-west europe showing the locations of earthquakes recorded from january 1970 to december 2004. earthquakes in denmark are from a catalogue at the geological survey of denmark and greenland, and the other earthquake locations are from a scandinavian catalogue at helsinki university. thick black lines show large postglacial faults c. 9000 years old, according to lagerbäck (1991). the curves in central scandinavia show an early common scandinavian compilation in project bifrost of uplift (mm/year). updated earthquake files for denmark are available at .geus.dk under seismology, and for the rest of scandinavia at www.seismo.helsinki.fi. the map is modified from gregersen & voss (2009). © 2012 geus. geological survey of denmark and greenland bulletin 26, 41–44. open access: www.geus.dk/publications/bull 4242 the last hundred years is even lower than that of northern scandinavia (fig.1). historical reports from the past thousand years tell the same story. the stresses in denmark are fairly homogeneous, mainly influenced by plate motion and only marginally by postglacial uplift, and the earthquake hazard is considered low (giardini 1999). here, we assess the longer geological perspective in an attempt to improve long term evaluation of hazard and to demonstrate the need for further geological investigations of stresses and strains in a time scale of several thousands of years. the area discussed is shown in fig. 2, which also shows generalised, relative sealevel changes after the time of the littorina sea in the midholocene (christensen 2001). discussion of selected areas carlsberg fault, copenhagen – this fault is one of the most significant faults in the copenhagen area (rosenkrantz 1937; nielsen & thybo 2004), and it has been very important for the carlsberg brewery because large quantities of water could be extracted from the fault zone for brewing beer. along the fault zone, cracks in house walls are observed, but it is an open question whether it is worse than in the adjacent area (ovesen et al. 2002; o.w. christensen, personal communication 1990s; b. larsen, personal communication 2009; l. nielsen, personal communication 2009). no earthquakes have been recorded near the fault. a more quantitative area mapping around the fault is needed. misfits in the danish geodetic system have been ascribed to the carlsberg fault. the base line of the distance measuring system on the island of amager near copenhagen has been described as deformed (ovesen et al. 2002). however, a more probable explanation of this so-called deformation is movements of near-surface sediments at one of the end points of the base line before 1911 due to winter freezing and thawing (m. aarestrup, personal communication 1991; k. engsager, personal communication 2009). we consider it unlikely that deformation occurred in 1930 as suggsted by rosenkrantz (1937) since no movements were registered between 1911 and 1933. another base line on amager (in danish called ‘prøvebane’, trial distance) has been reported as deformed (ovesen et al. 2002). in this case the geodetic problem was identified from observations which showed that there were problems with the end point as mentioned above. one of the ends is sloping strangely, and it must have been disturbed by winter freezing and thawing (k. engsager, personal communication 2009). we conclude that the carlsberg fault is not active at present. læsø – indications of a large fault with a displacement of 4 m have been reported from geological investigations of sand deposits younger than 7000 years on the island of læsø in kattegat (hansen 1977, 1980, 1994). no other sedimentary indications of earthquakes have been identified. hansen suggested that læsø does not fit into the post-glacial uplift pattern of the neighbouring coasts of kattegat, but according to christensen (2001) the elevations of the highest littorina sea shorelines on læsø fit well into the regional pattern for kattegat (fig. 2). a number of transgressions and regressions occurred over a period of several thousand years in the mid-holocene in the kattegat region (christensen 2001). the last of these so-called littorina transgressions ended around 4500 years ago. the structural differences in the sorgenfrei–tornquist zone tell about an older geological regime, not the present intraplate inactivity. hansen (1980) presented a curve showing changes of tilt of a succession of shorelines based on shoreline elevation measurements without any evaluation of uncertainties. each point in fig. 3 represents an average of 10–15 individual measurements that span an elevation of c. 1 m. this 1 m can be taken as a rough estimate of four times the standard error. if an estimated standard error of 0.25 m is taken into account, the argument for variations in tilting becomes nonsignificant. we find that the observed differences and the standard errors are of the same magnitude (fig. 3). the zigzag line of hansen (1980) goes through the average values of the slopes and the data may as well be explained by a straight line, i.e. within the belt of intervals determined by averages ± one standard error. hence we conclude that the differences hobro 50 km 10°e 10°e 56°n 8300 bc 8000 bc 7200 bc 6600 bc 0 2 6 8 10 12 4 læsø rügen skåne skagerrak kattegat sweden jylland 23 20 5600 bc 5600/4500 bc 6200 bc 6500 bc anholt copenhagen aalborg denmark amager fig. 2. map of south-western scandinavia, showing maximum elevations (m) and ages of shorelines from the littorina sea, according to christensen (2001). the circles show areas discussed in the text. 43 in average tilt cannot be used as an argument for discontinuous and occasionally reverse tilting. the data may as well agree with steady uplift and tilting. the same magnitude of standard errors in the use of geological markers for tilt measurements, and for elevation measurements with respect to sea level, is used in a recent work on anholt, also in the kattegat. when changes in sea currents and stormy weather which can influence the development of shorelines are added, the observations on shoreline tilting (fig. 3) cannot be used on their own as an argument for tectonic tilting. on the basis of our assessment of old and new arguments we conclude that the bedrock subsurface below læsø has been uplifted by the same amount as the rest of the kattegat region, regularly or with minor irregularities. this means that there was no læsø earthquake 4000–5000 years ago. the lower-lying parts of læsø with many well-documented beach ridges have their own exciting history (hansen et al. 2011), best explained by regular postglacial uplift together with coastal development influenced by sea currents, climate and weather. one hundred kilometres south-east of læsø, in the kattegat sea, where earthquakes have been registered (fig. 1), neotectonic small-scale faulting has been recognised (jensen et al. 2002). northern jylland – it has been suggested that a bend in the geodetic and geological uplift pattern north of aalborg may reflect a fault (lykke-andersen & borre 2000). however, according to follow-up measurements the bend could also be due to gradual creep with no indications of earthquakes (gregersen & schmidt 2001). central jylland – conspicuous so-called fracture valleys in central jylland were recently discussed by jakobsen & pedersen (2009). the valleys may be related to a deep fault zone, but no evidence exists for present-day earthquakes in this region. west coast of sweden – several localities on the west coast of sweden have been discussed by mörner (2003, 2009). signs of geological movements at these localities include variations in postglacial uplift of shorelines, faulting and rock deformation in a few cases, and in many incidences rock slides. also signs of liquefaction were found at several stratigraphic levels. we encourage a geological and geophysical assessment of these localities. skåne – in skåne in southern sweden repeated gps measurements have indicated differential movements of the two sides of the sorgenfrei–tornquist zone. however, these measurements are not considered statistically significant by swedish geodesists. skagerrak – in skagerrak and southern norway more earthquake activity takes place than in denmark (fig. 1). an assessment similar to the present one has been carried out by olesen et al. 2004. rügen – the bending phenomenon mentioned above is not confined to northern jylland. similar geodetic observations were reported from the island of rügen in northern germany (fig. 2; r. dietrich, personal communication 2008). however, seismological investigations show that northernmost germany is an earthquake-free region like southernmost denmark. discussion and conclusions we have evaluated and discussed a number of reports on geological or geodetic indications of earthquake activity in denmark. we found no signs of geologically recent faulting or recent crustal deformation. this corroborates that denmark and its neighbouring areas are characterised by a small earthquake potential. the largest earthquakes recorded had a magnitude of just below 6 on the richter scale. this is supported by the decrease of uplift stress after the last deglaciation. we conclude that earthquake hazard evaluations in denmark in the foreseeable future are best based on seismological data 200 100 0 –100 –200 c. 1 0 0 0 b c t ilt ( m m /k m t o w ar d s th e so u th -w es t) oldest succession of beaches youngest shoreline with one standard deviation fig. 3. tilt of shorelines on læsø (hansen 1980, fig. 5) with our estimated standard errors added. shorelines (dots) are plotted on the horizontal axis, with the youngest to the right. when the standard errors are included it appears that jumps in tilt values are not significant. the data can be interpreted as a steady change of tilt, approximated by the thick sloping line in the diagram, which is generally within one standard error and certainly within two standard errors. 4444 alone (gregersen & voss 2009). the earthquake zones are those found by historical and instrumental seismological investigations in the seas around denmark (fig. 1). unfortunately this does not mean that destructive earthquakes cannot happen. the situation in denmark is parallel to that of other intraplate regions. a totally unexpected destructive earthquake happens once every decade somewhere on earth within a quiet lithospheric plate. geodynamic stresses exist, which are occasionally released in structurally weak zones. based on our present state of knowledge of seismology, geodesy and geology, we cannot foresee in which fault zone or region an earthquake may happen in denmark, but no region is directly earthquake prone. not even the very significant sorgenfrei–tornquist zone is an earthquake zone (gregersen et al. 2011). acknowledgments we appreciate many good discussions with our colleagues jens morten hansen, peter johannesen, lars henrik nielsen, tine b. larsen, tanni j. abramovitz, abbas khan, karsten engsager, holger lykke-andersen, torben bidstrup, birger larsen, jørgen leth, lars nielsen, henrik olsen, hans thybo and lars b. clemmensen. references christensen, c. 2001: kystbosættelse og havniveauændringer i stenalderen. in: jensen, o.l., sørensen, s.a. & hansen, k.m. (eds): danmarks jægerstenalder – status og perspektiver, 183–193. hørsholm: hørsholm egns museum. giardini, d. (ed.) 1999: the global seismic hazard assessment program (gshap) 1992–1999. annali di geofisica 42, 280 pp. gregersen, s. & schmidt, k. 2001: tektonik i danmark. sorgenfrei–tornquist zonen. geologisk nyt 1, 16–17. aarhus: aarhus universitet. gregersen, s. & voss, p. 2009: stress change over short geological time: the case of scandinavia over 9,000 years since the ice age. in: reicherter, k., michetti, a.m. & silva barroso, p.g. (eds): palaeoseismology: historical and prehistorical records of earthquake ground effects for seismic hazard assessment. geological society special publications (london) 316, 173–178. gregersen s. & voss, p. 2010: irregularities in scandinavian postglacial uplift/subsidence in time scales tens, hundreds, thousands of years. journal of geodynamics 50, 27–31. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: sg@geus.dk gregersen, s., larsen, t.b. & voss, p. 2011: jordskælv i sorgenfrei–tornquist zonen? geoviden 1, 18–19. københavn: geocenter danmark. hansen, j.m. 1977: sedimentary history of the island læsø, denmark. bulletin of the geological society of denmark 26, 217–236. hansen, j.m. 1980: læsøs postglaciale udvikling i relation til den fennoskandiske randzone. dansk geologisk forening, årsskrift for 1979, 23–30. hansen, j.m. 1994: læsøs tilblivelse og landskaber – om øen der rokker og hopper, 55 pp. københavn: danmarks geologiske undersøgelse. hansen, j.m., aagaard, t. & binderup, m. 2011: absolute sea levels and isostatic changes of the eastern north sea to central baltic region during the last 900 years. boreas 41, 180–208. jakobsen, p.r. & pedersen s.a.s. 2009: fracture valleys in central jylland – a neotectonic feature. geological survey of denmark and greenland bulletin 17, 33–36. jensen, j.b., petersen, k.s., konradi, p., kuijpers, a., bennike, o., lemke, w. & endler, r. 2002: neotectonics, sea-level changes and biological evolution in the fennoscandian border zone of the southern kattegat sea. boreas 31, 133–150. khan, s., knudsen, p. & tscherning, c. 2005: crustal deformations at permanent gps sites in denmark. in: sansò, f. (ed.): a window on the future of geodesy. proceedings of the international association of geodesy symposia 128, 556–560. lagerbäck, r. 1991: seismically deformed sediments in the lansjärv area, northern sweden. skb technical report 91-17, 58 pp. stockholm: svensk kärnbränslehandtering ab. lykke-andersen, h. & borre, k. 2000: aktiv tektonik i danmark: der er liv i sorgenfrei–tornquist zonen. geologisk nyt 6, 12–13. aarhus: aarhus universitet. mörner, n.-a. 2003: paleoseismicity of sweden, a novel paradigm, 320 pp. stockholm: university of stockholm. mörner, n.-a. 2009: late holocene earthquake geology in sweden. in: reicherter, k., michetti, a.m. & silva barroso, p.g. (eds): palaeoseismology: historical and prehistorical records of earthquake ground effects for seismic hazard assessment. geological society of special publications (london) 316, 179–188. nielsen, l. & thybo, h. 2004: location of the carlsberg fault zone from seismic controlled-source fan recordings. geophysical research letters 31, l07621, http://dx.doi.org/10.1029/2004gl019603 olesen, o., blikra, l.h., braathen, a., dehls, j.f., olsen, l., rise, l., roberts, d., riis, f., faleide, j.i. & anda, e. 2004: neotectonic deformation in norway and its implications: a review. norwegian journal of geology 84, 3–34. ovesen, n.k., blem, h., gregersen, s., møller, h.m.f. & frederiksen, j.k. 2002: recent terrain-movements in copenhagen. dansk geoteknisk forenings bulletin 19, 183–192. rosenkrantz, a. 1937: bemærkninger om det østsjællandske daniens stratigrafi og tektonik. meddelelse fra dansk geologisk forening 9, 199–212. geological survey of denmark and greenland bulletin 20, 2010, 39–42 39 radon (222rn) is a radioactive, noble insoluble gas with a half-life of 3.8 days. it belongs to the uranium (238u) decay chain where radon is formed from radium (226ra). uranium and radium are built into mineral structures or are, for example, adsorbed on the surface of clay minerals, limonite or organic material. when radon is formed by radioactive decay from radium, parts of it enter the pores of rocks and soils and are transported by diffusive or advective forces in the pores. the transport rate depends on the permeability and water content in the pores (nazaroff 1992). radon may enter into buildings through fractures in the walls driven by forces such as pressure gradients between the outside and the inside. radon and its radioactive decay products are inhaled by living beings and are the main source of radiation to which humans are exposed. the radiation presents an increased risk of lung cancer and may cause leukaemia, which mainly occurs in children. on the basis of risk analyses it has been estimated that 10% of all cases of lung cancer in denmark are caused by radon inhalation (sundhedsstyrelsen 1987). it has been demonstrated that rocks and soils around and below houses are the main sources of radon emanation. several studies have analysed and described the radon content in danish sediments and rocks (e.g. damkjær & korsbech 1985; gravesen et al. 1996) and have demonstrated its relationship to radon levels in danish buildings (andersen et al. 2006; raaschou-nielsen et al. 2008). a nation-wide mapping of radon levels in danish dwelling houses based on, e.g. mapping of quaternary surface deposits and information about radon in sediments and rocks was performed by andersen et al. (2001). this paper presents some results concerning the radon content and emanation rates in different danish till deposits of saalian and weichselian age from a study carried out by the geological survey of denmark and greenland (geus). radon content in danish till deposits: relationship with redox conditions and age peter gravesen and peter roll jakobsen lundeborg viby bælthav re-advance east jylland ice border line main stationary line 50 km weichselian deposits saalian deposits weichselian outwash plains holocene marine deposits holocene aeolian deposits holocene freshwater deposits andrup thisted late glacial marine deposits © geus, 2010. geological survey of denmark and greenland bulletin 20, 39–42. open access: www.geus.dk/publications/bull fig. 1. simplified geological map of denmark, showing the general distribution of surface lithologies, the most important icemargin lines from the weichselian and the four investigated localities. 4040 geological setting clayey and sandy tills have been analysed for uranium, radium and radon content at the four localities viby, lundeborg, thisted and andrup (fig. 1). the andrup locality is situated outside the limit of the weichselian glaciation and the till is considered to be of saalian age deposited by a warthe ice advance about 180 000–160 000 years ago (houmark-nielsen 2007). the thisted, lundeborg and viby localities are located inside the maximum extent of weichselian glaciacion and the tills are referred to the weichselian. a relatively thin layer of till covers limestone and chalk deposits at thisted. the till is regarded as deposited during the late weichselian by the norwegian advance or from the main glacial advance from the north-east about 25  000–20  000 years ago (gry 1979). at lundeborg and viby the latest glacial advances are the east jylland advance and the bælthav re-advance about 19 000 years ago. in sections close to the lundeborg locality the slightly older mid danish till, deposited during the north-east advance, is seen below the till from the young baltic advance. some of the borehole samples may represent the mid danish till (houmark-nielsen & kjær 2003). materials and methods trenches were excavated at the four localities down to a depth of c. 2 m into till deposits (fig. 2). the sections were described with respect to lithology, structures and macropores, and samples of the clayey and sandy till units were collected at each locality. five 5–6 m deep boreholes were drilled with an 8 inch twist auger close to the excavations. the nearly undisturbed till samples were described, and samples were collected at 30 cm intervals. the samples from the excavations had a weight of at least 500–1000 g, and those from the boreholes 300–500 g. a total of 155 samples were collected, with each sample consisting of two subsamples. one subsample was collected in a sealed plastic bin for radon analysis, and the other was analysed for chemical compounds including uranium and radium. uranium was measured by instrumental neutron activation analysis at activation laboratories ltd and radium by a germanium detector at the national institute of radiation protection. the radon emanation rate was measured by the closed-chamber method using zns(ag) scintillation cells at risø national laboratory, danish technical university. prior to the radon analyses the samples were fragmented. the purpose was to increase the surface area of the sediment to promote the release of radon produced by radium decay. in addition to the chemical analyses, grain-size distribution, clast composition, organic content and water content were analysed at geus immediately after sampling. results and discussion this paper focuses on the measured radon emanation rates related to the weathering and redox (reduction–oxidation) conditions in the tills and the age difference between the tills at the four localities. till composition. the investigated tills are heterogeneous sediments with a matrix consisting of up to 20% clay, c. 70% silt and sand, and 4–10% gravel and stones. the clasts normally consist of norwegian and swedish basement rocks, quartz, danish chert and limestone fragments. the matrix is a mixture of crushed basement rocks and local sediments dominated by quartz, feldspar, mica, smectite, kaolinite and illite. as a consequence of the heterogeneous sediment types uranium and radium are related to the source material and not uniformly distributed. weathering and redox conditions. the upper 0–3 m of the till profiles have been subject to weathering after the last deglaciation (weichselian: thisted, lundeborg and viby; saalian: andrup). rainwater has percolated through pores in the till and caused decomposition and leaching of clay minerals, limonite and caco3. measurements of the caco3 content show deep weathering at andrup and nearly none at viby. most till profiles in denmark have an upper, yellow-brown coloured, oxidised zone and a lower, olive-grey-reduced zone separated by a redox interface. at andrup the redox interface coincides with the caco3 leaching boundary, while at lundeborg it is situated in the oxidised zone. at viby only slight caco3 leaching has occurred in the top of the profile. the redox zones are seen in the sections at andrup, lunderborg and viby. in table 1 the measured uranium and radium content and radon emanation rates are shown in relation to their occurrence in the two redox zones. the development fig. 2. field investigations at the lundeborg locality. drilling close to the excavated trench. 41 of the redox zones depends on chemical and physical processes and therefore the behaviour of u and ra in the zones is related to these processes. in the oxidised zone oxygen and for example nitrate influence the compositions of the minerals. the original uranium contained in the minerals is oxidised and mobilised, whereas the decay product radium is less mobile. radium is not uniformly distributed and often precipitated as films or crusts at the rim of the sediment pores between grains of clay minerals, limonite and caco3 where it produces radon. in the reduced zone the u-bearing minerals are more stable and emit less radon (ball et al. 1991). the investigation demonstrates larger rates of radon emanation in the oxidised zone than in the reduced zone (table 1). the radium and radon levels seem to be comparable with the measurements reported by damkjær & korsbech (1985). the radium concentration is fairly constant in the two redox zones (table 1). the higher radon content in the oxidised zone may be caused by enhanced emanation conditions. it may be due to opening of mineral pores in the oxidised zone as a result of an increase in number and size of macro-pores and higher permeability than in the reduced zone where radon is often trapped in smaller pores. viby sandy till weichselian 1.4–2.6 22.4–24.5 7.9–11.3 clayey till ~18 ka 17.6–26.1 22.1–35.9 5.2–13.9 4.5–6.7 lundeborg sandy till weichselian 1.1–1.9 18.5 7.3 clayey till ~18 ka 13.8–24.9 14.8–15.5 7.3–15.6 3.2–6.8 thisted sandy till weichselian 0.5–1.6 18.7–24.7 7.5–10.4 clayey till ~23 ka 14.1 13.6 limey till 1.2–6.6 0.9–1.5 andrup clayey till saalian 1.2–1.9 14.1–20.6 12.6–19.1 6.2–10.6 2.6–3.7 ~170 ka locality lithology age uranium radium radium radon radon (ppm) (oxidised) (reduced) (oxidised) (reduced) bq/kg bq/kg atom/kg/s atom/kg/s table 1. uranium and radium contents and radon emanation rates at the four localities 75 samples were analysed in total viby lundeborg thisted andrup radium (bq/kg) r ad o n ( u n it a to m s/ se c/ k g) oxidised till reduced till radium (bq/kg) r ad o n ( u n it a to m s/ se c/ k g) 0 5 10 15 20 25 30 35 40 0 5 10 15 20 25 30 35 40 0 5 10 15 20 25 30 35 40 0 5 10 15 20 25 30 35 40 a b fig. 3. comparison between radium content and radon emanation rates in danish tills. a radium content of 1 bq kg–1 yields a total radon production rate of 1 atom kg–1 sec–1. this relationship is represented by the diagonal line in the diagrams. a: plot of radium versus radon based on values for oxidised and reduced zones at the four studied localities. b: plot of radium versus radon in the oxidised and reduced zones at the four localities. one sample seems to produce more radon than the radium content can explain. the value is 15 616 ± 1710 atoms kg–1 sec–1. 4242 water content is also a factor to consider because it influences the distance the radon atoms have to travel to enter the pores. according to damkjær & korsbech (1985) up to 17% of water in the pores will increase the emanation rate. the water content in the tills is mostly below 17% but above the ideal content of 3–7%, and the content is slightly higher in the oxidised than in the reduced zone. radon emanation rates in the two redox zones appear to be independent of the geographical distribution of the till. in fig. 3 the distribution of radon and radium is seen in relation to redox zones. the radon emanation rates in both zones are lower than the corresponding radium content, which indicates that only some of the radon leaves the mineral grains and enters the pores of the till. the reduced zone has lower radon emanation rates than the oxidised zone in nearly all cases. transport conditions. after the radon atoms have moved into the pores the transport is controlled by the porosity and permeability of the till. till has a relatively high matrix porosity but the occurrence of macro-pores in the form of roots, fractures and faults is of considerable importance. macropores are most frequent in the oxidised zone. they were found in all sections, and their distribution explains the large difference in hydraulic conductivity between the two zones. the oxidised zone shows systematically decreasing values of hydraulic conductivity from the surface to a depth of 2–4 m (the redox interface) from 10–3 m/s to 10–11 m/s. below the redox interface the hydraulic conductivity varies unsystematically from 10–10m/s to 10–5 m/s (nilsson & klint 2009). age relationship. a comparison between radon and radium values from the saalian locality (andrup) and the weichselian localities (thisted, lundeborg and viby) show differences that appear to be related to the age of the till. figure 3 shows that average radon emanation rates generally increase from the saalian till to the younger weichselian till and it is seen that there is a clear difference between the andrup and viby localities. this is probably because the till at andrup has been exposed to weathering and precipitation for a longer time than that at viby. these processes have led to removal of uranium and fixed radium at a lower level than seen for the weichselian till. the in-situ radon producing radioactive decay in the saalian till lasted more than 150 000 years longer than in the weichselian till. another explanation could be different bulk mineralogy in the tills deposited by ice advances of different ages. final remarks our investigation indicates that radon emanation rates in till are higher in the oxidised zone than in the reduced zone at all four localities and that they are higher in weichselian till than in saalian till. however, only limited data are available for this conclusion. many danish dwelling houses are built on clayey tills, and the new information concerning the redox conditions is important when evaluating transport and impact of radon into dwelling houses and the risk of exposure to inhalation of radon in denmark. acknowledgement the ministry of health and prevention is thanked for financial support. references andersen, c.e., ulbak, k., damkjær, a. & gravesen, p. 2001: radon i danske boliger. kortlægning af lands-, amtsog kommuneværdier, 132 pp. copenhagen: sundhedstyrelsen, statens institut for strålehygiejne. andersen, c.e., raaschou-nielsen, o., andersen, h.p., lind, m., gravesen, p., thomsen, b. & ulbak k. 2006: prediction of 222rn in danish dwellings using geology and house construction information from central databases. radiation protection dosimetry 27, 10–21. ball, t.k., cameron, d.g., colman, t.b. & roberts, p.d. 1991: behaviour of radon in the geological environment: a review. quarterly journal of engineering geology and hydrogeology 24, 169–182. damkjær, a. & korsbech, u. 1985: measurement of the emanation of radon-222 from danish soils. the science of the total environment 45, 343–350. gravesen, p., jakobsen, p.r. & kelstrup, n. 1996: radon i danske jordarter ii. undersøgelser og konklusioner. danmarks og grønlands geologiske undersøgelse rapport 1996/78, 113 pp. gry, h. 1979: beskrivelse til geologisk kort over danmark. kortbladet løgstør. kvartære aflejringer 1:100  000/1:50  000. danmarks gelogiske undersøgelse i. række 26, 58 pp. houmark-nielsen, m. 2007: extent and age of middle and late pleistocene glaciations and periglacial episodes in southern jylland, denmark. bulletin of the geological society of denmark 55, 9–35. houmark-nielsen, m. & kjær, k. 2003: southwest scandinavia, 40–15 kyr bp: palaeogeography and environmental change. journal of quaternary science 18, 769–786. nazaroff, w.w. 1992: radon transport from soil to air. reviews of geophysics 30, 137–160. nilsson, b. & klint, k.e. 2009: bilag 7. sammenstilling af data fra hydrauliske undersøgelser i moræneler – videnstatus. in: gravesen, p. & rosenberg, p. (eds): særligt pesticidfølsomme lerområder: datagrundlag og mulige veje mod zonering, kupa, 70 pp. + appendices (report). copenhagen: geological survey of denmark and greenland. raaschou-nielsen, o., andersen, c.e., andersen, h.p., gravesen, p., lind, m., schüz, j. & ulbak, k. 2008: domestic radon and childhood cancer in denmark. epidemiology 19, 536–543. sundhedsstyrelsen 1987: radon-boliger-strålingsdosis-lungekræftrisiko, 14 pp. copenhagen: statens institut for strålehygiejne. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: pg@geus.dk geological survey of denmark and greenland bulletin 23, 2011, 61–64 61 scattered occurrences of palaeogene sediments are found in north-east greenland, where they overlie unconformably cretaceous sediments and are capped by palaeogene basalts. these sediments have received little attention (watt 1994), except for relatively recent studies (nøhr-hansen & piasecki 2002; jolley & whitham 2004; larsen et al. 2005; heilmann-clausen et al. 2008). as part of an ongoing petroleum geological study that focuses on the jurassic–cretaceous succession, the palaeogene sediments were included to better constrain their age, depositional environment and relation to the basalts. several localities were investigated on wollaston forland, sabine ø and hold with hope, a few of which are described here (fig. 1). eastern wollaston forland and sabine ø discontinuous outcrops of mostly loose and un-cemented palaeogene sediments occur in haredal, eastern wollaston forland. a n–s-striking normal fault with 100–125 m of downthrow to the east separates the main outcrop in the southern slope into two blocks. the best exposed succession is situated in the western footwall block, where the succession dips 20° to the sw (fig. 2). it overlies marine mudstones of late albian age (wigginsiella grandstandica subzone (v1) of nøhr-hansen 1993) in the footwall block and of early to middle campanian age (indicated by the dinocysts alterbidinium ioannidesii and cerodinium diebelii) in the hanging wall block; however, the contact to the cretaceous is not exposed. a poorly exposed palaeogene succession in the northern slope is probably from the hanging wall block; the base of the succession and underlying strata are not exposed. haredal, southern slope. approximately 200 m of palaeogene sediments are partly exposed in the footwall block, forming two upward-coarsening units overlain by basalts (fig. 2). the lower unit consists of more than 70 m of dark grey mud overlain by 54 m of fineto medium-grained sand beds alternating with thinner heteroliths topped by more than 2 m of coarse-grained sand. the succession is of earliest ypresian age based on the presence of the dinocyst apectodinium augustum (fig. 3) and the nannofossil discoaster lenticularis. the palynological assemblage is dominated by reworked material from upper jurassic (e.g. gonyaulacysta jurassica), mid to upper cretaceous (e.g. hapsocysta benteae, chatangiella spp. and wodehouseia spinata) and lower paleocene (e.g. alisocysta margarita). the presence of a. augustum may correlate with the a. augustum (p6b) subzone described from the central north sea (mudge & bujak 1996) and correlated palaeogene deposits in north-east greenland henrik nøhr-hansen, lars henrik nielsen, emma sheldon, jussi hovikoski and peter alsen fig. 1. geological map of the wollaston forland – hold with hope study area in north-east greenland. the location of the studied sections corresponds to the distribution of palaeogene sediments. © geus, 2011. geological survey of denmark and greenland bulletin 23, 61–64. open access: www.geus.dk/publications/bull young sund g a e l h a m k e b u g t lille pendulum kap broer ruys sabine ø wollaston forland hold with hope palnatoke bjerg ice quaternary intrusive complexes (48–28 ma) basaltic sills and dykes basaltic plateau lavas (58–54 ma) palaeogene cretaceous jurassic triassic palaeozoic daneborg tobias dal 20°w 74°n jackson ø harefjeld haredal fosdalen langsiden dronning augusta dal 25 km home forland faults studied section greenland 6262 e le va ti o n 400 m 380 360 340 320 300 280 260 240 220 200 e . p al ae o ge n e p e ri o d /e p o ch st ag e palyno zones palyno zones c h ro n o st ra ti gr ap h y c h ro n o st ra ti gr ap h y p 6 a . au gu st um z o n e p 6 b a . au gu st um su b zo n e sa m p le s pa la eo p er id in iu m p yr op ho ru m a p ec to di ni um a ug us tu m a re ol ig er a sp p . c er od in iu m s tr ia tu m d ef la nd re a oe bi sf el de ns is pa la eo cy st od in iu m b ul lif or m e t ha la ss ip ho ra d el ic at a a p ec to di ni um p an ic ul at um g la p hy ro cy st a re tii nt ex ta a lis oc ys ta m ar ga ri ta c er od in iu m d ie be lii pa la eo cy st od in iu m a us tr al in um c er od in iu m s p ec io su m ph th an op er id in iu m s p p . ? ? ?? ? ? ? ? ? ? ? ?? ? ? ? ? ? ? dinoflagellate cysts haredal southern slope, eastern wollaston forland haredal southern slope haredal northern sloped is co as te r le nt ic ul ar is d is co as te r ge m m eu s? h el ic os p ha er a se m in ul um t ri br ac hi at us o rt ho st yl us na n p z o n e s z o n e e le va ti o n 220 a. augustum top a. augustum base a. augustum base a. margarita base t. evittii top m 200 180 160 140 120 100 80 60 e . p al ae o ge n e p e ri o d /e p o ch st ag e p 6 a . a ug us tu m z o n e su b zo n e sa m p le s a re ol ig er a sp p . c er od in iu m d ie be lii g la p hy ro cy st a cf . re tii nt ex ta pa la eo p er id in iu m p yr op ho ru m sp on go di ni um d el iti en se t ri th yr od in iu m e vi tt ii se no ni as p ha er a in or na ta a lis oc ys ta m ar ga ri ta c or do sp ha er id iu m in od es d in o cy st s p 4 h n h 2 0 0 2 m ur at od in iu m f im br ia tu m pa la eo cy st od in iu m b ul lif or m e se ne ga lin iu m it er la ae ns e c er od in iu m s tr ia tu m a p ec to di ni um a ug us tu m a p ec to di ni um p an ic ul at um d ef la nd re a oe bi sf el de ns is t ha la ss ip ho ra d el ic at a a p ec to di ni um h om om or p hu m ? ? ? ? ? ? ? ? dinoflagellate cysts e c re t. l a lb . (v) (v1) p 6 b a . au gu st um n p 9 n p 1 0 n p 1 1 u . n p 1 0 n p 1 1 ? n p 1 2 clay silt sand ? ? ? ? pebbl.clay silt gnl. sand cross bedding cross lamination wavy and flaser bedding erosive contact clay and coal clast sandstone mudstone muddy sandstone basalt conglomerate lithology planar lamination sedimentary structures ? 220 ? ? ? g ?? l .. t h an et ia n ? e . y p re si an l . t h an et ia n ? e . y p re si an ? ? 1 2 3 4 a. augustum top e . d an ia n ? e . se la n d ia n a b fault west east with the paleocene–eocene thermal maximum (petm) which occurred at about 56 ma and lasted for c. 170 kyr (harding et al. 2011). the upper, coarser-grained part only yielded two dinocyst species cerodinium sp. and areoligera sp. and no nannofossils. the upper unit consists of more than 20 m of dark grey silty and sandy mud overlain by slightly heterolithic, finegrained sand (22 m), fine-grained sand (10 m), conglomerate and pebbly sandstone (8 m) topped by a few metres of sand, mud and carbonaceous mud or coal. the palynologifig. 2. a: southern slope and fault at haredal. the highest mountain is c. 700 m high. b: sedimentological logs and range charts of selected in situ dinocysts and calcareous nannofossils. numerals in the zones of the northern slope are: 1: alisocysta margarita zone, 2: senegalinium iterlaaense zone or palaeocystodinium bulliforme zone, 3: trithyrodinium evittii zone and 4: spongodinium delitiense subzone (nøhr-hansen et al. 2002). na: nannofossils, np: palaeogene nannoplankton zone. v, v1: subtilisphaera kalaalliti zone and wigginsiella grandstandica subzone (nøhr-hansen 1993). 63 cal assemblages below the conglomerate are dominated by reworked material of mid – late cretaceous age and a few specimens of paleocene age. the succession is of early ypresian (early eocene) age based on poorly preserved specimens of the dinocysts apectodinum augustum, apectodinium paniculatum, deflandrea oebisfeldensis and the nannofossil tribrachiatus orthostylus (fig. 3). samples above the conglomerate yielded a few reworked upper cretaceous and paleocene dinocysts together with a few indeterminate algae and the nannofossil helicosphaera seminulum (fig. 3). the latter indicates an age not younger than mid ypresian for the youngest dated sediments in haredal; this is compatable with a mid ypresian radiometric 39ar/40ar age of 55.02 ± 0.49 ma for the oldest lava analysed from wollaston forland (l.m. larsen personal communication 2008). haredal northern slope. approximately 180 m of sediments were studied on the northern slope of haredal (fig 1). the lower 40 m consist of dark grey mud with a palynological assemblage dominated by reworked material from mid – upper cretaceous strata (e.g. hapsocysta benteae, chatangiella spp. and aquilapollenites spp.). the presence of trithyrodinium evittii, spongodinium delitiense and a few specimens of senoniasphaera inornata in the two lowermost samples indicates an early paleocene age and may correlate with the lower danian trithyrodinium evittii zone (fig. 2) established from west greenland (nøhr-hansen et al. 2002). the presence of common senegalinium iterlaaense and palaeocystodinium bulliforme in the two overlying samples indicates a mid to late danian age correlating with the senegalinium iterlaaense and palaeocystodinium bulliforme zones (2 and 3 in fig. 2), whereas the next sample contains a few specimens of the dinocysts alisocysta margarita and cerodinium striatum indicating a late danian/?early selandian age correlating with the alisocysta margarita zone (nøhr-hansen et al. 2002; 1 in fig. 2). the upper 130 m of sand and mud contain very few in situ palynomorphs, however the presence of apectodinium augustum indicates correlation with the apectodinium augustum (p6b) subzone, indicating an ?early selandian – thanetian hiatus. outer haredal, dronning augusta dal and sabine ø. palaeogene sediments from the southern slope of the easternmost part of haredal, on the eastern slope of dronning augusta dal, and on the north-eastern slope of harebjerg, sabine ø were also studied and sampled (fig. 1).the successions all contain the petm dinocyst marker apectodinium augustum. north-eastern hold with hope langsiden. interbedded in dark grey mud at langsiden (fig. 1) occurs an 8 m thick unit with sharp-based, upwardfining successions of conglomerate and pebbly sand, up to a few metres thick containing large reworked mudstone clasts. the palynological assemblage from the underlying mud is dominated by a reworked flora of late maastrichtian age indicated by the presence of triblastula wilsonii and wodehouseia octospina. the presence of in situ trithyrodinium evittii, spongodinium delitiense and a few specimens of senoniasphaera inornata indicates correlation with the lower danian trithyrodinium evittii zone (nøhr-hansen et al. 2002). the mudstone clasts yielded a mid cretaceous flora. the palynomorph assemblage of two samples from the overlying mud is likewise dominated by a reworked upper cretaceous flora; however, the lower sample also contains a few specimens of the dinocysts alisocysta margarita and cerodinium striatum, indicating correlation with the upper danian/?lower selandian alisocysta margarita zone (nøhrhansen et al. 2002). the presence of thalassiphora delicata in the upper sample indicates a latest danian/early selandian age (nøhr-hansen & piasecki 2002). east of fosdalen. approximately 6 m of loose, white-grey, fineto medium-grained sand with scattered small clay clasts 1 42a 2b 3 65 987 10 20 μm 20 μm 20 μm 20 μm 20 μm 20 μm 20 μm 2 μm2 μm2 μm20 μm fig. 3. images of selected dinocysts (1–7) and nannofossils (8–10). 1: trithyrodinium evittii. 2a, b: alisocysta margarita. 3: thalassiphora delicata. 4: gen et sp. indet. of piasecki et al. (1992). 5: apectodinium augustum. 6: apectodinium paniculatum. 7: deflandrea oebisfeldensis. 8: discoaster gemmeus? 9: helicosphaera seminulum. 10: tribrachiatus orthostylus. 6464 overlain by a less than 2 m thick bed of dark grey mud covered by volcanic rocks overlies mid cretaceous sandy mudstone east of fosdalen (fig. 1). the palynological assemblage of the mud bed is dominated by spores and pollen and some reworked dinocysts of mid to late cretaceous age. in situ specimens of the dinocyst gen. et sp. indet. of piasecki et al. (1992; fig. 3) also occur, suggesting fresh to brackish water. the species is common in wells offshore eastern canada just above the apectodinium augustum p6b subzone (h. nøhrhansen, unpublished data). the occurrence of the species immediately below the basalts may indicate an earliest ypresian age. discussion the new biostratigraphic dating shows that the palaeogene sediments on wollaston forland, hold with hope and sabine ø comprise paleocene and earliest eocene strata with a hiatus that probably spans the major part of the selandian and thanetian. the age of the underlying cretaceous deposits east of the fault in haredal is early–middle campanian, much younger than the middle albian previously described from wollaston forland (nøhr-hansen 1993). the stratigraphic gap between the cretaceous and palaeogene sediments thus decreases towards the basin to the east. the large amount of reworked cretaceous marine palynomorphs including a late maastrichtian flora documents uplift of cretaceous marine sediments and major erosion during the early palaeogene. the ages of the youngest palaeogene sediments and the oldest flood basalts appear to be almost identical; however, the nature of the contact between the sediments and basalts needs to be further investigated to determine whether the contact is angular, as is suggested in places by a relatively steep dip of the sediments. discontinuous conglomerate beds with rounded quartzite pebbles and boulders up to 20 cm, as well as cross-bedded sandstones, terminate the upper unit in haredal and are interpreted as fluvial channel deposits. the absence of basaltic clasts indicates that deposition occurred prior to the volcanic events. the thin overlying succession with a coaly bed is interpreted as an aggrading coastal plain subject to marine inundations. however, the principal part of the palaeogene sediments accumulated in a marine environment as indicated by marine dinocysts and the presence of scattered marine trace fossils. thin, fine-grained sandstone beds with flute casts and a massive lower part overlain by beds with parallel lamination and cross-lamination in outer haredal indicate deposition from turbidite currents. the sharp-based, finingupward conglomerate beds and pebbly sand from langsiden embedded in mud are interpreted as channelised gravity flow deposits. the two upward-coarsening units in haredal suggest that the palaeogene sediments mainly accumulated during two major depositional phases. potential by-pass surfaces are identified at the cretaceous–palaeogene boundary, at the top of the lower unit, at the base of the fluvial conglomerates and possibly at the sediment–basalt boundary. coarsegrained sediments may have been transported toward the basin area to the east along these surfaces. references harding, i.c. et al. 2011: sea-level and salinity fluctuations during the paleocene–eocene thermal maximum in arctic spitsbergen. earth and planetary science letters 303, 97–107. heilmann-clausen, c., abrahamsen, n., larsen, m., piasecki, s. & stemmerik, l. 2008: age of the youngest paleogene flood basalts in east greenland. newsletters on stratigraphy 43, 55–63. jolley, d.w. & whitham, a.g. 2004: a stratigraphical and palaeoenvironmental analysis of the sub-basaltic palaeogene sediments of east greenland. petroleum geoscience 10, 53–60. larsen, m., heilmann-clausen, c., piasecki, s. & stemmerik, l. 2005: at the edge of a new ocean: post-volcanic evolution of the palaeogene kap dalton group, east greenland. in: doré, a.g. & vining, b. (eds): petroleum geology: north-west europe and global perspectives. proceedings of the 6th petroleum conference london 2, 923–932. london: geological society. mudge, d.c. & bujak, j.p. 1996: paleocene biostratigraphy and sequence stratigraphy of the uk central north sea. marine and petroleum geology 13, 295–312. nøhr-hansen, h. 1993: dinoflagellate cyst stratigraphy of the barremian to albian, lower cretaceous, north-east greenland. grønlands geologiske undersøgelse bulletin 166, 171 pp. nøhr-hansen, h, & piasecki, s. 2002: palaeocene age of sub-basaltic sediments at savoia halvø, east greenland. geology of greenland survey bulletin 191, 111–116. nøhr-hansen, h., sheldon, e. & dam, g. 2002: a new biostratigraphic scheme for the paleocene onshore west greenland and its implications for the timing of the pre-volcanic evolution. in: jolley, d.w. & bell, b.r. (eds): the north atlantic igneous province: stratigraphy, tectonic, volcanic and magmatic processes. geological society, special publications (london) 197, 111–156. piasecki, s., larsen, l.m., pedersen, a.k. & pedersen, g.k. 1992: palynostratigraphy of the lower tertiary volcanics and marine clastic sediments in the southern part of the west greenland basin: implications for the timing and duration of the volcanism. rapport grønlands geologiske undersøgelse 154, 13–31. watt, w.s. 1994: stratigraphy and correlation of the tertiary plateau basalts in north-east greenland. rapport grønlands geologiske undersøgelse 162, 185–194. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: hnh@geus.dk geological survey of denmark and greenland bulletin 23, 2011, 9–12 9 the cambrian to lower silurian succession in denmark is mostly composed of organic-rich black shales that were deposited in an epicontinental sea during a period of high global sea level (haq & schutter 2008). the mid-cambrian to early ordovician alum shale was intensively studied in the 1980s for its source-rock properties (e.g. buchardt et al. 1986). recent attention has focused on its potential as an unconventional shale gas source (energistyrelsen 2010). on southern bornholm, many wells have been drilled through the lower palaeozoic succession because of its importance for groundwater exploitation. in western denmark, only the deep exploration wells slagelse-1 and terne-1 have penetrated the alum shale, and knowledge of the unit west of bornholm is thus very limited (fig. 1). the project ‘shale gas in europe (gash)’ was launched in 2009 to address the european shale gas potential (horsfield et al. 2008) and is organised by the german research centre of geosciences and sponsored by oil and energy companies. it deals with basic research of key aspects of gas shale from regional to reservoir scales, and focuses on four ‘natural laboratories’, namely the lower palaeozoic alum shale, the carboniferous (namurian) shales, the lower jurassic posedonia shale and as a reference the north american barnet shale. as part of gash, a european black shale database is under construction in order to facilitate exploration and exploitation of gas shales in europe. the geological survey of denmark and greenland (geus) has taken part in both the gas shale research and in the national data compilation for the european shale database. this paper presents the results of drilling on bornholm in august 2010 by geus with the aim of obtaining fresh core material relevant to shale gas studies within the gash project (the skelbro-2 core) and providing new stratigraphic and geochemical information on the lower palaeozoic (the billegrav-2 core). in particular, logging of the silurian was needed to improve the log-stratigraphical template of pedersen & klitten (1990) which will enable the correlation of geophysical logs from non-cored water wells. drilling and logging project on bornholm the southern part of bornholm is characterised by a mosaic of fault blocks (graversen 2010). outcrops, old core data and logs from water wells were used to find the best drilling locations for the two new wells. both wells were fully cored and subsequently subjected to an extensive logging programme by geus in order to characterise the lithology as well as the water composition and the flow capacity of the fracture systems (fig. 2). spectral gamma and density scanning of the cores was subsequently carried out at geus. skelbro-2 well the skelbro-2 well was drilled 275 m east of skelbro-1 (dgu 246.749; pedersen 1989). the mid-ordovician komstad limestone is 4 m thick at this locality. the top of the alum shale was encountered at 8.5 m below surface, and a total of 33.5 m of alum shale was drilled. the well was terminated at 42.9 m in the lower cambrian rispebjerg member (læså formation). the cored succession is virtually identical to the one described by pedersen (1989) from the skelbro-1 core. billegrav-2 100 km skelbro-2 caledonian front slagelse-1 terne-1 bornholm norwegian–danish basin ringkøbing–fyn high fig. 1. map of denmark showing the distribution of lower palaeozoic strata (modified from buchardt et al. 1997) and the location of the slagelse-1 and terne-1 exploration wells, and the skelbro-2 (dgu 246.817) and billegrav-2 (dgu 248.61) wells. shale gas investigations in denmark: lower palaeozoic shales on bornholm niels hemmingsen schovsbo, arne thorshøj nielsen, kurt klitten, anders mathiesen and per rasmussen © geus, 2011. geological survey of denmark and greenland bulletin 23, 9–12. open access: www.geus.dk/publications/bull 1010 l .c . m .c . u p p er o rd o vi ci an l o w er s ilu ri an fu ro n gi an l .o . q grey shaleblack shale sandstonelimestone siltstone (no log below 90 m)(no log below 90 m) k lithology depthperiod (m) natural gamma (api)10 formation resistivity (ohm-m)10 fluid conductivity (mscm–1)600 flow log (%)0 10 20 30 40 50 60 70 80 90 100 110 120 1600800 100 strat. r .m . a lu m s h al e fm d ic el lo gr ap tu s sh . l in d eg år d d1 d2 d3 a b1 b2 b3 b4 f3 f5 f4 e1 e2 e3 f1 f2 log unit r as tr it es s h al e (api)10 2000 250 sonic velocity (km/s) 62 fig. 2. stratigraphy, lithology, natural gamma ray, formation resistivity, sonic velocity, fluid conductivity and flow logs for billegrav-2. log units a–f according to pedersen & klitten (1990). dashed lines: uncertain biostratigraphical boundaries. green arrows: major inflow zones; black arrows: minor inflow zones. q: quaternary. k: komstad limestone. l.o.: lower ordovician. l.c.: lower cambrian. m.c.: middle cambrian. r.m.: rispebjerg member. api: american petroleum institute (a standard unit for gamma-ray measurements). note change in scale for natural gamma at 90 m. 11 billegrav-2 well the billegrav-2 well was drilled close to locality 14b of bjerreskov (1975) and 800 m south of the billegrav-1 well (dgu 247.560; pedersen 1989). the silurian rastrites shale was cored from 4.5 m below surface down to 60.5 m. the rastrites shale comprises light to dark mudstone except for a distinct grey mudto siltstone unit containing carbonatecemented sandy beds at a depth between 31.2 and 46.0 m (fig. 2). the upper ordovician includes the lindegård formation (previously referred to as the tretaspis shale or tommarp and jerrestad mudstones), comprising grey mudand siltstone and the dark organic-rich dicellograptus shale (fig. 2). the base of the dicellograptus shale is located at a depth of 95 m. in its lowermost part the shale contains numerous bentonite beds including a 1 m thick k-bentonite bed that represents an important regional marker bed (bergström & nilsson 1974). the komstad limestone is 0.1 m thick and only represented by its basal conglomerate. a thin bentonite rests directly on the komstad limestone conglomerate; there is no conglomerate at the base of the overlying dicellograptus shale. the alum shale formation is 27 m thick and includes the middle cambrian andrarum and exsulans limestone beds that are important regional marker beds (nielsen & schovsbo 2006). the base of the alum shale was reached at a depth of 122 m, and the well was terminated at 125.9 m in the rispebjerg member (fig. 2). log-stratigraphy of the billegrav-2 well correlation of water wells based on gamma variation has served as an effective mean of correlation between wells (pedersen & klitten 1990). all gamma-ray, log-defined units identified in nearby water wells can also be recognised in the billegrav-2 well (fig. 2). the resistivity and sonic logs provide important additional information (fig. 2). in the rastrites shale, the resistivity is particularly powerful in resolving the lithological variation, since the carbonate-cemented sandy beds in the middle part (the f3 unit) stand out as high-resistivity beds (fig. 2). water-flow zones in the billegrav-2 well the flow log from this well clearly indicates that the water flow is related to three major and three minor influx zones (fig. 2). the majority of the water flow takes place in the lowermost part of the rastrites shale at around a depth of 56–50 m. relatively high water influx is also seen in the rastrites shale at about a depth of 20 and 30 m, whereas there is very little water flow from log unit f3 (fig. 2). water influx is seen in the dicellograptus shale at 77 and 90 m as well as in the uppermost part of the alum shale at 96 m. no flow is observed deeper in the well. interestingly, the conductivity data suggest that the pore water in the alum shale has a much higher conductivity than in the shales above, suggesting that it is stagnant. this indicates that the alum shale acts as a hydraulic barrier between the siltand sandstone aquifers below and the shale aquifers above this unit. potential shale gas units onshore denmark shale gas units of potential economic interest have to be (1) matured to at least the gas generative stage, (2) organic rich (total organic carbon (toc) >2 wt%), (3) volumetrically important (thickness >20 m and regionally distributed) and (4) preferentially located away from structurally complex areas. the lower palaeozoic succession on bornholm contains up to 10 wt% toc in the alum shale, up to 5 wt% in the dicellograptus shale and up to 2 wt% in the rastrites shale (buchardt et al. 1986). also mesozoic organic-rich units occur in denmark, but those onshore are all thermally immature to marginally mature (petersen et al. 2008) and thus have no potential for shale gas. thermal maturity and burial of the lower palaeozoic succession the thermal maturity of lower palaeozoic shales in denmark is only known from a few wells that all have vitrinite reflectance values >2.5%, indicative of a post-mature rank with regard to oil generation (fig. 3). the shales are thus theoretically favourable for shale gas. the high maturity reflects deep burial and a high geothermal gradient in late silurian – early devonian time (buchardt et al. 1997). the present burial depth of the lower palaeozoic can be evaluated from the pre-zechstein depth map (fig. 4) that represents the deepest level that can be mapped with some config. 3. maturity of the lower palaeozoic sequence based on reflectance of vitrinite-like particles. the thermal maturity increases towards the caledonian front, reflecting deep burial in late silurian to early devonian time. modified from buchardt et al. (1997). ? ? ? ? ? ? immature mature over-mature anchi-metamorphic caledonian front 1212 fidence on a regional basis (vejbæk 1997). lower palaeozoic shales are buried to very deep levels in the central parts of the norwegian–danish basin and are probably not within reach of shale gas exploration. lower palaeozoic shales buried to moderate depths of 2–4 km occur in a broad belt around the margin of the norwegian–danish basin. shale gas investigations have focused on this region (energistyrelsen 2010). conclusions shale gas research and exploration in denmark is currently focused on palaeozoic shales. the prospective units are poorly known in denmark outside bornholm. the play involves deeply buried, post-mature shale in which many basic rock properties are still unknown. key questions that remain to be addressed include the gas storage capacity of the shales, their mineralogy and how they respond to fracturing. the hydrogeology of water wells on bornholm may provide a test case to assist in unravelling how fracture systems and flow are distributed in the lower palaeozoic shales. acknowledgements the gash project financed the skelbro-2 drilling. funding for the billegrav-2 drilling was provided by geus, the natural history museum of denmark, the university of southern denmark and bornholm’s regionskommune. peter and kristian turner from faxe kalk a/s assisted with the drilling. the land owners andres ipsen, jørn erik koefoed and jesper koefoed kindly permitted us to drill on their properties. references bergström, s.m. & nilsson, r. 1974: age and correlation of the middle ordovician bentonites on bornholm. bulletin of the geological society of denmark 23, 27–48. bjerreskov, m. 1975: llandoverian and wenlockian graptolites from bornholm. fossils and strata 8, 1–94. buchardt, b., clausen, j. & thomsen, e. 1986: carbon isotope composition of lower palaeozoic kerogen: effects of maturation. organic geochemistry 10, 127–134. buchardt, b., nielsen, a.t. & schovsbo, n.h. 1997: alun skiferen i skandinavien. geologisk tidsskrift 3, 1–30. energistyrelsen 2010: denmark’s oil and gas production 2009, 156 pp. copenhagen: the danish energy agency. graversen, o. 2010: structural analysis of superposed fault systems of the bornholm horst block, tornquist zone, denmark. bulletin of the geological society of denmark 57, 25–49. haq, b.u. & schutter, s.r. 2008: a chronology of paleozoic sea-level changes. science 322, 64–68. horsfield, b., schulz h.-m. & gash team 2008: gash: a shale gas initiative for europe. egu general assembly. geophysical research abstracts 10, egu 2008-a-01508. nielsen, a.t. & schovsbo, n.h. 2006: cambrian to basal ordovician lithostratigraphy in southern scandinavia. bulletin of the geological society of denmark 53, 47–92. pedersen, g.k. 1989: the sedimentology of lower palaeozoic black shales from the shallow wells skelbro 1 and billegrav 1, bornholm, denmark. bulletin of the geological society of denmark 37, 151–173. pedersen, g.k. & klitten, k. 1990: anvendelse af gamma-logs ved korrelation af marine skifre i vandforsyningsboringer på bornholm. danmarks geologisk forening årskrift 1987–89, 21–35. petersen, h.i., nielsen, l.h., bojesen-koefoed, j.a., mathiesen, a., kristensen, l. & dalhoff, f. 2008: evaluation of the quality, thermal maturity and distribution of potential source rocks in the danish part of the norwegian–danish basin. geological survey of denmark and greenland bulletin 16, 66 pp. vejbæk, o.v. 1997: dybe strukturer i danske sedimentære bassiner. geologisk tidsskrift 4, 31 pp. fig. 4. depth to the top pre-zechstein surface. note the depocentre in the norwegian–danish basin (>5000 m) and the shallow ringkøbing–fyn high (<1000 m). areas where the top pre-zechstein surface coincides with basement are shown in grey. modified from vejbæk (1997). faults at top pre-zechstein top pre-zechstein on basement 5000 1000 depth to top pre-zechstein (m) 100 km authors’ addresses n.h.s., k.k., a.m. & p.r., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: nsc@geus.dk a.t.n., natural history museum of denmark, øster voldgade 5–7, dk-1350 copenhagen k, denmark. geological survey of denmark and greenland bulletin 42, 2018, 7-14 7 the upper jurassic blokelv-1 cored borehole in jameson land, east greenland – an introduction morten bjerager, stefan piasecki and jørgen a. bojesen-koefoed the geological survey of denmark and greenland (geus) successfully drilled the fully cored blokelv-1 borehole in the central part of the jameson land basin in east greenland, targeting the upper jurassic, rich source-rock interval of the hareelv formation. the borehole achieved 100% core recovery from 1.72 m to a total depth of 233.8 m; the recovered hareelv formation section consists of interlayered black, laminated organic-rich mudstones, massive sandstones and heterolithic sandstone–mudstone intervals of the katedralen member, and amalgamated massive sandstones of the sjællandselv member. the core is of very high quality and has been subjected to an extensive sampling and analytical programme focused particularly on petroleum geological aspects, as presented in the following eight papers in this volume. this bulletin describes an important, previously poorly documented member of the ‘kimmeridge clay’ family of prolific petroleum source rocks in the north atlantic area. keywords: east greenland, jameson land, upper jurassic, hareelv formation, cored borehole ___________________________________________________________________________ geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. email: mbj@geus.dk the blokelv-1 borehole was drilled by the geological survey of denmark and greenland (geus) in jameson land during the summer of 2008. it was the first of a series of fully cored boreholes planned as part of a collaborative project between geus and a number of sponsoring oil companies entitled petroleum geological studies, services and data in east and north-east greenland (bojesenkoefoed et al. 2009, 2014). the borehole is situated in the central, deep part of the jurassic jameson land basin near the blokelv river, after which the borehole is named, c. 35 km west of the constable pynt airport (fig. 1). the borehole targeted a poorly exposed part of the upper jurassic hareelv formation that represents a midkimmeridgian sea-level highstand (surlyk 2003) and is a correlative of the well-known kimmeridge clay formation (sensu lato) of northwest europe; it was thus expected to include a rich petroleum source-rock interval. the north-east greenland part of the circum arctic resource appraisal (cara; christiansen et al. 2006; gautier 2007; gautier et al. 2011) pointed out that the quality of an upper jurassic – lower cretaceous source rock in north-east greenland should be considered a major risk element in assessing the petroleum potential of the region. data based on outcrop sampling, however, have shown surprisingly low petroleum potential. a drilling programme that included the blokelv-1 borehole was thus designed to provide as complete coverage as possible of the entire oxfordian to ryazanian succession onshore east and north-east greenland in order to better understand the geological context of the source rocks and to provide fresh and unweathered samples for new and comprehensive analyses of the entire succession. the succession cored by the blokelv-1 borehole covers the interval from the middle oxfordian to the lower volgian and in general terms the analytical programme included: • detailed sedimentological description and analysis, including interpretation of depositional environments represented by the deposits © geus, 2018. geological survey of denmark and greenland bulletin 42, 7–14. available at: www.geus.dk/bulletin42 mailto:mbj@geus.dk http://www.geus.dk/bulletin42 88 • detailed biostratigraphic analysis, based on both palynomorphs (dinoflagellate cysts) and macrofossils (ammonites) • detailed analysis of the petroleum potential of mudstones, including richness, thermal maturity, biomarker and stable carbon isotopic fingerprints of source rocks and oil stains • detailed analysis of the reservoir properties of sandstones, including diagenetic studies • chemostratigraphy based on bulk geochemical analysis and provenance analyses of sandstones based on dating of detrital zircons • analysis of uplift history based on apatite fission track analysis (afta) • analysis of magmatic intrusions penetrated. the results of the analytical programme have been previously reported to the group of sponsoring oil companies (bjerager et al. 2009). following the expiration of the 5-year confidentiality clause on 1 january 2015, the data were released for exclusive use by geus, and this bulletin presents the key results in a series of papers, each dealing with different aspects of the analyses. drilling operations the drilling operation was organised by geus and comprised a camp of six persons, including two geologists, two technicians, one of whom also undertook catering, and two drillers. the drilling was part of a major fieldfig. 1. geological map of the southern part of jameson land showing the location of the blokelv-1 borehole. only named river systems are indicated. ; ; ; ; ; ; jyllandselv blo kelv falst erelv lolla ndse lv ran unk ele lv ra uk elv gø01_02_273_03_mbj_blokelv_intro h urry in le t 24°w 23°w 22°w katedralen fortet jameson land liverpool land ugleelv gåseelv hareelv sjællandselv 20 km olympen fm fossilbjerget fm / pelion fm neill klinter gp kap stewart gp triassic basement hareelv fm, katedralen mb major dyke/sill ice quaternary hesteelv fm raukelv fm hareelv fm, salix dal mb hareelv fm, sjællandselv mb ?devonian 71°n 70°30'n mbj intro fig 1 s c o r e s b y s u n d blokelv-1 9 work programme, and a helicopter was chartered for the full season. two days of reconnaissance prior to drilling ensured selection of the optimal location for the drill site, remote from exposed major sills and dykes. the disassembled drill-rig parts were flown as sling loads from the constable pynt airport to the blokelv-1 drill site at 70°45.305´n / 23°40.430´w, at an elevation of 181 m above sea level (fig. 1). the drill rig was assembled directly from the helicopter sling (fig. 2). mobilisation of the drill rig and camp took three days, with six tons transported in helicopter sling and cabin loads. the blokelv-1 borehole was initiated (‘spudded’) on 7 august 2008 and completed on 15 august 2008 at a total depth (td) of 233.8 m. drilling was carried out using a diamant boart 747 wireline rig from faxe kalk a/s, with a casing diameter of 85 mm and 3 m long core barrels, yielding high-quality core material with a diameter of 56 mm. core recovery was 100% (1.72–233.8 m). the borehole was abandoned as an open hole with casing down to 20 m. the casing was cemented and closed at the top and a small cairn of sandstone blocks was built over the wellhead. wires with temperature sensors were a b fig. 2. transport and erection of the faxe kalk a/s diamant boart 747 wireline rig; the rig was transported as helicopter sling loads (a) and the drill tower was assembled directly (b). fig. 3. aerial view of the drill camp and rig; the well-trodden path leads from the mess tent to the drill site where the water basins are conspicuous. 1010 installed in the hole down to 184 m. demobilisation and down-hole logging took two days. drilling demanded a daily water consumption of 10 m3, with 80 m3 used in total. water was supplied from the nearest river at about 300 m lateral distance and 75 m vertical distance; this was achieved using a petrol-driven water pump and inflatable hoses to supply two 5000-litre basins (fig. 3). in addition, a 3000 litre basin of saline water was used to circulate through the borehole at night thereby avoiding freezing of the drill string in contact with the permafrost. permafrost was present from a few decimetres below the surface down to about 100 m. weather conditions were characteristic of the region being predominantly calm and sunny with temperatures up to c. 20°c during the day, interrupted by a few days of fog or low clouds and one day with strong winds. drilled succession the low plateaux in the terrain around the blokelv-1 drill site consist of sandstone of the upper jurassic (lower volgian) sjællandselv member underlain by poorly exposed alternating black mudstone and grey – yellowish grey sandstones of the katedralen member of the hareelv formation (figs 4, 5). the core section has been dated by biostratigraphy, the zonation being based on combined ammonite and dinoflagellate data (alsen & piasecki 2018, this volume). the upper part (0–10.08 m) is assigned to the sjællandselv member of the hareelv formation and consists of fractured and structureless, marine gravity-flow sandstones with a few thin very dark grey mudstone beds that contain fossils indicating the lowermost volgian p. elegans chronozone. the interval from 10.08 m to td at 233.8 m is assigned to the katedralen member of the hareelv formation, spanning the middle oxfordian to lowermost volgian c. densiplicatum – p. elegans chronozones. the cored katedralen member comprises very dark grey to black laminated mudstones (54%), sandstone/mudstone heteroliths (12%), and weakly stratified – structureless or intrusive sandstones (34%) (figs 6, 7). the base of the hareelv formation was not reached in the cored interval according to the lithological characteristics, but the lowermost part of the recovered core is assigned to the c. tenuiserratum and c. densiplicatum chronozones, which is time equivalent to the upper part of the olympen formation (larsen & surlyk 2003; bruhn & surlyk 2004). possible thin bentonite beds are present at 25.8 m, 79.9 m, 83.2 m and 216 m. palaeogene intrusions are represented by three sills (0.7 m, 1.2 m and 1.9 m thick at depths of about 27, 56 and 102 m, respectively) and a dyke (0.3 m thick at c. 7 m depth). logging a full wireline logging programme was planned to include conductivity, spectral gamma and sonic logs. the borehole was unstable at certain intervals, however, probably due to drilling-induced fractures in homogeneous sandstone. in-hole logging therefore only included fig. 4. outcropping sandstones of the sjællandselv member, hareelv formation near the drill site. fig. 5. poorly exposed mudstones of the hareelv formation near the drill site, emphasising the importance of obtaining fresh rock samples in cores for petroleum geological analysis. 11 chronostratigraphy m id dl e up pe r lo w er up pe r vo lgi an o xf or di an ki m m er id gia n u pp er ju ra ss ic lo w er grtotal cps conductivity ms/m0 400 30 000 40 000 10 depth (m) 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 210 220 230 10 depth (m) 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 210 220 230 233.8 be be gø04_02m_453_02_mbj_blokelv_intro mbj intro fig 6 mudstone heterolith (mudstone/sandstone) sandstone sandstone, remobilised (intruded) lithology igneous intrusion be ripple cross-lamination wavy bedding slump sandstone intrusion small mudstone intraclasts bentonite large mudstone clast coalified wood belemnite ammonite bivalve brachiopod structures, biota parallel lamination/bedding diffuse stratification clay si sand pebbl. fig. 6. simplified lithological log of the blokelv-1 core showing the conductivity log and gamma-ray log measured in the field. note that the wireline gamma-ray log was limited to the lowermost c. 45 m due to technical problems; these data were subsequently supplemented by a core spectral gamma-ray log. note the marked conductivity response of the three thin igneous intrusive bodies. cps: counts per second. ms/m: millisiemens per metre. 1212 top 106.3 m box 29 top 229.66 m box 62 top 184.07 m base 109.85 m base 233.40 m base 187.95 m box 50 fig. 7. core photographs of boxes 29, 50, and 62 showing black organic-rich mudstone, mudstone conglomerate, massive sandstone, laminated heterolithic mudstone–sandstone, slumped heterolithic sandstone–mudstone and sandstone dykes in the upper jurassic katedralen member (hareelv formation). the individual core sections are 1 m long. 13 a complete conductivity log and a partial gamma log (232–184 m depth; fig. 6). to compensate for this, a high-resolution spectral gamma log and density log of the complete core were later measured in the core laboratory at geus (bjerager et al. 2018a, this volume). the temperature in the borehole was measured 25.5 and 40.5 hours after drilling ceased at depths of c. 10 m, 35 m, 85 m and 184 m. after 40.5 hours, the temperature stabilised and the estimated depth of the base of the permafrost is at 100 m (table 1). wire and sensors were left in the borehole for later measurements when temperature conditions are in full equilibrium. sampling and analytical programme at the drill site, 79 full core samples for gas analysis were collected immediately from the bottom of each recovered core (spacing of 3 m) and stored in sealed metal cans. in addition, eight samples were collected for preliminary and ‘express’ biostratigraphic studies immediately after the drilling operation. three samples of observed bitumen were collected in the cores, including liquid bitumen internally in a belemnite, and solid bitumen along fractures (bojesen-koefoed et al. 2018, this volume). core photographing was conducted under moist and surface-dry conditions. the core was sedimentologically logged at a scale of 1:20 for detailed facies analysis and re-drafted at scales of 1:50 and 1:500 for overview logs (bjerager et al. 2018a, this volume). an extensive sampling programme was subsequently conducted in the laboratories at geus. biostratigraphic studies are based on 50 mudstone samples for palynology and 42 levels with ammonites (alsen & piasecki 2018, this volume). chemostratigraphic and diagenetic studies are based on 42 samples for bulk chemistry, 27 thinsections for petrography, and 22 xrd samples for bulk sample and clay mineralogy (bjerager et al. 2018a, this volume; olivarius et al. 2018a, this volume). provenance studies are based on six samples for identification of heavy minerals and zircon ages (olivarius et al. 2018b, this volume). reservoir properties were evaluated on 25 plugs and comprise measurements on grain density, porosity and permeability (bjerager et al. 2018a, this volume). petroleum geological evaluation is based on 139 samples for source rock quality and organic maturity with 42 samples for biomarker analysis, 22 samples for stable carbon isotope analyses (bojesen-koefoed et al. 2018, this volume). two samples of igneous intrusions were subjected to major and trace element analysis (larsen 2018, this volume). uplift studies were focused on two samples for apatite fission-track analysis (afta; green & japsen 2018, this volume). acknowledgements logistical assistance from polog ( j. weiss andersen) and constable pynt airport is gratefully acknowledged. faxe kalk a/s drillers peter turner and andy milton skilfully managed the drilling and together with john boserup and annette ryge ran an excellent drilling camp. an as350 helicopter chartered from air greenland transported the drilling equipment from constable pynt to the blokelv-1 drill site. the drilling project was carried out under the auspices of the petroleum geological studies, services and data in east and north-east greenland collaboration agreement. illustrations were prepared by jette halskov and stefan sølberg. references alsen, p. & piasecki, s. 2018: biostratigraphy of the hareelv formation (upper jurassic) in the blokelv-1 core, jameson land, central east greenland. in: ineson, j. & bojesen-koefoed, j.a. (eds): petroleum geology of the upper jurassic – lower cretaceous of east and north-east greenland: blokelv-1 borehole, jameson land basin. geological survey of denmark and greenland bulletin 42, 15–37 (this volume). bjerager, m., alsen, p., bojesen-koefoed, j., kjøller, c., larsen l.m., nytoft, h.p., olivarius, m., petersen h.i., piasecki, s. & schovsbo, n. 2009: blokelv corewell, ggu511101, upper jurassic hareelv formation in jameson land, east greenland. danmarks og grønlands geologiske undersøgelse rapport 2009/86, 3 volumes, 223 pp., 8 appendices. bjerager, m., alsen, p. bojesen-koefoed, j.a., piasecki, s. & pilgaard, a. 2018b: late jurassic evolution of the jameson land basin, east greenland – implications of the blokelv-1 borehole. in: ineson, j. & bojesen-koefoed, j.a. (eds): petroleum geology of the upper jurassic – lower cretaceous of east and north-east greenland: blokelv-1 borehole, jameson land basin. geological survey of depth/m temperature temperature after 25.5 hours after 40.5 hours table 1. temperature in the blokelv corehole, post drilling 0 m 12.8°c 8.5°c 10 m 12.5°c 0°c 35 m 12.0°c -1.5°c 85 m 13.0°c -1.0°c 184 m 12.7°c 2.2°c 1414 denmark and greenland bulletin 42, 149–168 (this volume). bjerager, m., kjøller, c., olivarius, m., olsen, d. & schovsbo, n. 2018a: sedimentology, geochemistry and reservoir properties of upper jurassic deep marine sediments (hareelv formation) in the blokelv-1 borehole, jameson land basin, east greenland. in: ineson, j. & bojesen-koefoed, j.a. (eds): petroleum geology of the upper jurassic – lower cretaceous of east and north-east greenland: blokelv-1 borehole, jameson land basin. geological survey of denmark and greenland bulletin 42, 39–64 (this volume). bojesen-koefoed, j.a., bjerager, m. & piasecki, s. 2009: shallow core drilling and petroleum geology related field work in north-east greenland 2008. geological survey of denmark and greenland bulletin 17, 53–56. bojesen-koefoed, j.a., alsen, p. & christiansen, f.g. 2014: six years of petroleum geological activities in north-east greenland (2008– 2013): projects and a view of the future. geological survey of denmark and greenland bulletin 31, 59–62. bojesen-koefoed, j.a., peter nytoft, h.p., petersen, h.i., piasecki, s. & pilgaard, a. 2018: petroleum potential of the upper jurassic hareelv formation, jameson land, east greenland. in: ineson, j. & bojesen-koefoed, j.a. 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(eds): petroleum geology of the upper jurassic – lower cretaceous of east and north-east greenland: blokelv-1 borehole, jameson land basin. geological survey of denmark and greenland bulletin 42, 133–147 (this volume). larsen l.m. 2018: igneous intrusions in the cored upper jurassic succession of the blokelv-1 core, jameson land basin, east greenland. in: ineson, j & bojesen-koefoed, j.a. (eds): petroleum geology of the upper jurassic – lower cretaceous of east and north-east greenland: blokelv-1 borehole, jameson land basin. geological survey of denmark and greenland bulletin 42, 133–147 (this volume). larsen, m. & surlyk, f. 2003: shelf-edge delta and slope deposition in the upper callovian – middle oxfordian olympen formation, east greenland. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 931–948. olivarius, m., bjerager, m., knudsen, c., keulen, n., & kokfelt, t. 2018b: provenance of basinal sandstones in the upper jurassic hareelv formation, jameson land basin, east greenland. in: ineson, j. & bojesen-koefoed, j.a. (eds): petroleum geology of the upper jurassic – lower cretaceous of east and north-east greenland: blokelv-1 borehole, jameson land basin. geological survey of denmark and greenland bulletin 42, 115–126 (this volume). olivarius, m., weibel, r., schovsbo, n.h., olsen, d. & kjøller, c. 2018a: diagenesis of upper jurassic sandstones of the blokelv-1 core in the jameson land basin, east greenland. in: ineson, j & bojesen-koefoed, j.a. (eds): petroleum geology of the upper jurassic – lower cretaceous of east and north-east greenland: blokelv-1 borehole, jameson land basin. geological survey of denmark and greenland bulletin 42, 65–84 (this volume). surlyk, f. 2003: the jurassic of east greenland: thermal subsidence, onset and culmination of rifting. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 659–722. surlyk, f., gjelberg, j. & noe-nygaard, n. 2007: the upper jurassic hareelv formation of east greenland: a giant sedimentary injection complex. in: hurst, a. & cartwright, j. (eds): sand injectites: implications for hydrocarbon exploration and production. aapg memoir 87, 141–149. _________________________________________________________________________________________ manuscript received 26 october 2015; revision accepted 26 february 2018 geological survey of denmark and greenland bulletin 23, 2011, 33-36 33 in the summer of 2010, the geological survey of denmark and greenland (geus) mapped the potential raw materials and substrate types, over large parts of the danish economic sector of the north sea, in cooperation with orbicon a/s. the mapping was carried out for the danish nature agency; it is part of the general mapping of raw material resources within the territories of the danish state and forms part of the input for the implementation of the european union’s marine strategy framework directive. the purpose was (1) to provide an overview of the distribution, volume and composition of available raw materials and (2) to identify, describe and map the distribution of the dominant marine bottom types. methods during the first part of the field work a single beam echo sounder for bathymetrical data, a side scan sonar for mapping the seabed surface as well as a chirp (1–10 khz) and a sparker (1 khz) were used to map the layers below the seabed with up to 50 m penetration. after preliminary interpretation, the acoustic data acquisition was followed by sediment sampling using a vibrocorer with up to 6 m penetration and a grab sampler in order to investigate the characteristics of the sediments. in addition, the seabed was filmed using a video camera mounted on a remotely operated vehicle (rov) to document the seabed sediment substrate and benthic fauna. the field work was divided into three phases (fig. 1, table 1): phase 1: mapping of both raw materials and bottom types, phase 2: mapping of bottom types alone and phase 3: geological mapping. bathymetry and geological model the data from the phase 1 area made it possible to map the bathymetry and develop a geological model for the area. the bathymetric data show a nw–se-trending ridge 25–30 m below present sea level interpreted as the offshore continuation of the main stationary line (msl) that formed during the last glacial maximum (fig. 2a; houmark-nielsen & kjær 2003). the high area comprises the core areas of jyske rev and lille fisker banke. north of the msl a series of ridges trending nne–ssw that are 20–40 km long and 5–10 km broad dominate the bathymetry. the crests of the ridges are 18–24 m b.s.l. and the troughs around 40 m b.s.l. these large ridges are interpreted as early holocene giant tidal sand ridges. backstripping shows the pre-holocene transgression surface and accentuates the msl (fig. 2b). south of the msl the bathymetry shows a gentle south-western slope, which we interpret as drowned sandur deposits that formed in front of the melting scandinavian ice sheet. mapping of raw materials and habitats in the danish sector of the north sea jørn bo jensen, sara borre, jørgen o. leth, zyad al-hamdani and laura g. addington 50 km 4°e 6°e 56°n 57°n 8°e jylland border of the exclusive economic zone vibrocore video grab sample and video phase 1 seismic grid phase 2 seismic grid phase 3 seismic grid a b c main stationary line (msl) fig. 1. map of the economic sector of the danish north sea showing the seismic grids and sampling positions in the combined raw material and marine bottom-type mapping carried out in 2010. a: phase 1 area. b: phase 2 area. c: phase 3 area. © geus, 2011. geological survey of denmark and greenland bulletin 23, 33–36. open access: www.geus.dk/publications/bull 3434 the pre-quaternary surface is close to the seabed in the phase 1 area (figs 3, 4; nielsen et al. 2008) where it is overlain by a few tens of metres of quaternary sediments. thicker quaternary deposits are found in local depressions connected with salt domes and in n–s-trending buried tunnel valleys that are incised into the pre-quaternary deposits (huuse & lykke-andersen 2000). the quaternary sediments overlie miocene sand and clay, danian limestone and locally cretaceous chalk. the quaternary stratigraphy in the north sea is poorly known, but south of the msl elsterian and saalian till deposits have been identified on the seismic sections and in sediment cores in a situation similar to that found in the onshore hill islands in western jylland. to some extent the deeply incised valleys are filled with pre-weichselian sand and eemian marine silt and sand (figs 3, 4). to a large extent the weichselian deposits are related to the msl. they are partly seen as glaciotectonic deformations in the sparker data, and partly as till deposited below and at the margin of the ice sheet. glaciofluvial sand and gravel are located in depressions in the weichselian landscape and proximal to distal sandur sediments overlying glacial deposits south of the msl (figs 3, 4). late glacial marine deposits overlie the glacial deposits at present water depths greater than 40 m, which reflects the relative sea level around 17 000 years bp (leth 1996), shortly after the last deglaciation (fig. 3). from c. 17 000 years bp the relative sea level dropped to reach a minimum of about 50 m b.s.l. at c. 11 000 years bp. at that time only the northernmost part of the phase 1 area was below sea level (fig. 2b) and a widespread hiatus is seen that lasted until submergence in connection with the holocene transgression. the relative sea level rose continuously from c. 11 000 to c. 6000 years bp to reach a maximum level at 1–3 m above the present sea level. over the past 6000 years the relative sea level fell to its present level. the holocene deposits can be divided into four units: 50–52 48–50 46–48 44–46 42–44 40–42 38–40 36–38 34–36 32–34 30–32 28–30 26–28 24–26 22–24 20–22 18–20 depth (m) 20 km 7°e 7°e 57°n 56°30´n 7°e 7°e a b jyskejyske revrev lille fiskerlille fisker bankebanke jyske rev lille fisker banke msl msl phase survey grid (km) acoustic profiles (km) vibrocores grab samples video data 1 2 × 5 c. 3000 60 24 85 2 7 × 7* / 7 × 15 c. 3000 – 31 67 3 7 × 7 c. 2500 – – – * eastern area, western area. table 1. details of phases 1–3 # # fig. 2. bathymetry and palaeo-morphology of the phase 1 area (fig. 1), off north-west jylland. a: bathymetry. b: palaeo-morphology in metres below present sea level of the pre-holocene transgression surface. the weichselian main stationary line (msl) is located in the lille fisker banke – jyske rev area, between the stippled lines. 35 early holocene giant tidal sand ridge deposits. the giant tidal sand ridges are connected to the msl which trends nne–ssw and attains heights up to 20 m (figs 2–4). the architecture implies that the ridges were formerly connected to the shore and formed by tidal currents during the holocene sea-level rise, mostly by trapping of sediment within tidal eddies generated by headlands or flow convergence. similar giant sand ridges related to holocene sea-level rise have been reported from the english channel (reynaud et al. 2003). early holocene fine-grained marine and fjord sediments. incised valleys formed in the easternmost part of the phase 1 area during the lowstand period and during the initial part of the holocene transgression (fig. 3). fine-grained marine sediments were deposited in such valleys in protected areas east of jyske rev, and brackish fine-grained sedimentation took place in palaeo-fjords similar to the present limfjorden (agger clay; leth 1996). holocene beach ridge and spit deposits. these accumulations formed during the progressive transgression when the glacial deposits were isolated as islands in the open sea. erosion, transport and deposition led to the formation of beach deposits on the lee side of jyske rev. sub-recent to recent mobile sand. around 8000–7000 years ago the shallow parts (18–20 m b.s.l.) of the phase 1 area were transgressed by the sea. the present wave system and the jutland current developed (gyllencreutz et al. 2006), which resulted in the formation of mobile sand waves and major sand banks in the jyske rev and lille fisker banke region. the mobile sand deposits reached several metres in thickness and the bedforms can have wavelengths of over 100 m. potential for raw material we have identified a number of stratigraphical units that could be of interest to the sand and gravel industry. preweichselian sand in buried tunnel valleys is of potential interest in areas where no or little cover sediment is found. pre-quaternary deposits pre-weichselian till pre-weichselian valley fill (sand?) eemian marine silt and sand weichselian till weichselian glaciofluvial sand and gravel weichselian sandur silt and sand late glacial marine silt and sand early holocene valley fill border of deep valley holocene marine sand and gravel 7°e 57°n 56°30´n 20 km 2000600010 00014 000 d ep th ( m ) age (years bp) -40 -20 0 fig. 3. seabed surface sediments in the phase 1 area (fig. 1). the location of the geological profile (fig. 4) is shown with a red stippled line. the inset diagram is a model of relative sea-level changes after the last deglaciation in the study area. nnw se d ep th ( m b el o w s ea l ev el ) 30 40 60 50 5 km sse nw fig. 4. nw–se geological profile through the phase 1 area (fig. 1) covered in the first phase. for legend and location of profile see fig. 3. 3636 this is seen in the south-eastern part of the phase 1 area, but little is known about the volume and quality of the material accumulated in the tunnel valleys. weichselian glaciofluvial sand and gravel deposits along the msl are of great interest as a resource for concrete production. however, this resource occurs only sporadically and is often covered by several metres of holocene sediments. giant holocene sand ridges in the north-western part of the phase 1 area are the most noteworthy sand resource in the region because of their enormous volume of around 8 × 109 m3. the resource is easily accessible and consists of mediumgrained sand that is well suited for land reclamation, beach nourishment and concrete production. drowned coastal deposits are the traditional dredged resource in the jyske rev area. this resource consists of high quality sand and gravel for concrete production as the high energy level during erosion and transport of the sediments has removed the light particles. recent mobile sand is often deposited in sand waves and the sediment consists of well-sorted, medium-grained sand that is excellent for beach nourishment. however, before this resource is removed it is crucial to ensure that the sand waves are not part of the present sand budget of the coast. geological seabed units as bottom types and possible sand-eel fishery areas bottom-type mapping was part of the seabed mapping task for the danish nature agency. since the seabed sediments form the habitat for marine benthic organisms and are of vital importance to the distribution of marine life, the seabed sediments have been used for mapping the dominant bottom types. a number of additional parameters such as light penetration, salinity and temperature influence the distribution of faunal types, but a close link between till deposits on jyske rev and the natura 2000 code 1170 stone reef type (boedeker et al. 2006) is obvious. a close link is also seen between giant sand ridges and the natura 2000 code 1110 sand bank type. the mapping of the geological unit giant sand ridges as sand banks was compared with the distribution of sand-eel fishing grounds in the jyske rev and lille fisker banke areas (jensen et al. 2011) and a nearly perfect match was found. concluding remarks we have developed a geological model that can form the basis of combined mapping of raw material and marine bottom types in the lille fisker banke and jyske rev areas. weichselian glaciofluvial deposits, early holocene giant, tidal sand ridges, middle holocene drowned, coastal sand and gravel deposits and sub-recent to recent mobile sand waves and banks form potential or substantiated raw material geological units. examples of dominant bottom types are the jyske rev till deposits that are classified as a stone reef of the seabed type, and the early holocene giant tidal sand ridges that represent a sand bank type. sand-eel fishing grounds have been used as an example of the close linkage between geology, bottom types and fish habitats. acknowledgement the danish nature agency is thanked for permission to publish the paper. references boedeker, d., krause, j.c. & von nordheim, h. 2006: interpretation, identification and ecological assessment of the natura 2000 habitats ‘sandbank’ and ‘reef ’. in: von nordheim, h., boedeker, d. & krause, j.c. (eds): progress in marine conservation in europe, natura 2000 sites in german offshore waters, 47–64. berlin: springer. gyllencreutz, r., backman, j., jakobsson, m., kissel, c. & arnold, e. 2006: postglacial paleoceanography in the skagerrak. the holocene 16, 975–985. houmark-nielsen, m. & kjær, k.h. 2003: southwest scandinavia 40–15 kyr bp: palaeogeography and environmental change. journal of quaternary science 18, 769–786. huuse, m. & lykke-andersen, h. 2000: overdeepened quaternary valleys in the eastern danish north sea: morphology and origin. quaternary science reviews 19, 1233–1253. jensen, h., rindorf, a., wright, p.j. & mosegaard, h. 2011: inferring the location and scale of mixing between habitat areas of lesser sandeel through information from the fishery. ices journal of marine science 68, 43–51. leth, j.o. 1996: late quaternary geological development of the jutland bank and the initiation of the jutland current, ne north sea. geological survey of norway bulletin 430, 25–34. nielsen, t., mathiesen, a. & bryde-auken, m. 2008: base quaternary in the danish parts of the north sea and skagerrak. geolological survey of denmark and greenland bulletin 15, 37–40. reynaud, j.-y., tessier, b., auffret, j.-p., berné, s., de batist, m., marsset, t. & walker, p. 2003: the offshore quaternary sediment bodies of the english channel and its western approaches. journal of quaternary science 18, 361–371. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: jbj@geus.dk geological survey of denmark and greenland bulletin 31, 2014, 51-54 51 a multidisciplinary study of a geothermal reservoir below thisted, denmark morten leth hjuler, henrik vosgerau, carsten møller nielsen, peter frykman, lars kristensen, anders mathiesen, torben bidstrup and lars henrik nielsen the first geothermal plant in denmark was established in 1984 near the town of thisted (fig. 1). for nearly 30 years the plant has successfully produced c. 43°c hot water (surface temperature) from a highly permeable sandstone reservoir in the late triassic to early jurassic gassum formation and used the heat from the geothermal water for district heating. the 45°c hot water (formation temperature) is pumped up from a vertical production well, thisted-2, from a depth of c. 1250 m and the cooled water (c. 12°c) is re-injected into the formation through a vertical injection well, thisted-3, located 1.5 km east of the production well. in order to increase the capacity of the plant the supplier of district heating, thisted varmeforsyning, plans to add a new well to the current configuration. in 2013 the geological survey of denmark and greenland was assigned the task to propose suitable sites for a new well. the sites will be proposed based on the quality, continuity and temperature of the reservoir(s) within the gassum formation. three possible well sites (thisted-5a–c) were considered (fig. 1). well-log information and core-analysis data from four existing wells, thisted-1–4, indicate significant variations in reservoir properties and spatial extent of the reservoir bodies within the gassum formation. in order to establish a robust geological reservoir model that covers all reservoir-qualifying aspects it was decided to conduct the evaluation as a multidisciplinary study that included: (1) seismic interpretation and mapping, (2) determination of reservoir temperature, (3) petrophysical interpretation of well logs and (4) sedimentological description of well cores. two possible well configurations were considered: (1) the new thisted-5 is used as an injection well and thisted-2 continues as a production well. production can be increased with additional injection capacity, and by using thisted-5 as the principal injection well the breakthrough of cold injection water is postponed. (2) the new thisted-5 is used as a production well with thisted-2 and -3 as injection wells. thisted-5 will be located downflank of the thisted salt structure and will produce geothermal water from greater depths and expectedly with higher temperature than the existing production well. geological setting and seismic mapping thisted is located in the central to northern part of the norwegian–danish basin in an area characterised by numerous salt pillows and diapirs. the upper permian – mesozoic suc© 2014 geus. geological survey of denmark and greenland bulletin 31, 51–54. open access: www.geus.dk/publications/bull possible site for new well contour line 1500 land sea town thisted existing well geothermal plant seismic line 1 km 1000 thisted-3 thisted-4 thisted-2 thisted-1 thisted-5c thisted-5b thisted-5a 750 1250 fig. 2 fig. 1. depth structure map of near top gassum formation (contour interval: 50 m) of the thisted area showing seismic lines, well locations and possible locations of the planned thisted-5 wells. base upper cretaceous mid-cimmerian unconformity near top gassum formation near base gassum formation near top skagerrak formation near top zechstein near base zechstein 0 1 2 3 2 km nw se tw ow ay tr av el ti m e (s ec ) fig. 2. seismic section (prkl7374a no 74249) across the thisted salt pillow. 5252 cession consists of 5–5.5 km of sediments (vejbæk & britze 1994). thick zechstein salt layers are overlain by triassic sandstone, mudstone, carbonate and salt, followed by lower jurassic mudstone, middle jurassic sandstone, upper jurassic – lower cretaceous mudstone and siltstone with few sandstone beds. this succession is overlain by thick layers of chalk and limestone. thisted is situated over the southern part of a gently sloping salt pillow and the strata above the salt pillow in the area of the thisted-2 and 3 wells dip towards the south. thisted-1 and 4 are located on top of the structure. all relevant 2d seismic profiles in the thisted area were interpreted in order to map the presence and variations of the reservoirs including depth, changes in thickness and the occurrence of faults which may affect the lateral continuity of the reservoirs. due to low resolution of the seismic data near the potential sites of thisted-5, it was difficult to map faults and lateral changes in lithology and to trace seismic horizons. however, seven horizons could be identified, including the near top and near base of the gassum formation (fig. 2). mapping of single reservoir intervals within the gassum formation was impossible and no faults were identified. based on the seismic interpretation, four depth structure maps were compiled, including one of the near top gassum formation (fig. 1). from this map the depth of the gassum formation at the three suggested well sites can be estimated. relative to the thisted-2 well, the thisted-5a site is located downflank of the thisted structure whereas the thisted5b and 5c are situated upflank. due to deeper burial and expected higher temperature of the geothermal target, the thisted-5a site is suitable as a production well, whereas thisted-5b and 5c are injection well candidates. reservoir temperature temperature data from danish onshore wells are limited and include values from different depths and different formations, thus the geothermal gradient covers a wide range of 28–20°c/km. in the thisted-2 production well, c. 43°c warm water (surface temperature) is produced from a depth of 1250 m (formation temperature 45°c). the geothermal target in thisted-5a is estimated to be located at a depth of c. 1450 m, which corresponds to a temperature up to 52°c, if extrapolated from a continuation of the regional gradient and the temperature data from the thisted-3 well. the geothermal target in the thisted-5b and 5c injection sites is located at a depth of c. 1150 m corresponding to a reservoir temperature of 41°c using thisted-3 data. 6 260 ms/ft dt 60 cali 16 prs g as su m f or m at io n fj er rit sle v fm vi nd in g fm 1130 1140 1150 1160 1170 1180 1190 1200 1210 1220 1230 1240 1250 1260 1270 1280 1290 1300 1310 1320 gr api0 250 inches phie 0 % 40110000 md perm_log 1 rhob 1.7 g/cm3 2.7 nphi 0.6 v/v dec 0 depth (m) sandstone siltstone shale marlstone potential reservoir sand production interval fig. 3. petrophysical evaluation of the gassum formation in the thisted-2 well, including a lithological interpretation. 0 1 10 100 1000 10 000 100 000 0 5 10 15 20 25 30 35 40 45 50 ai r pe rm ea bi lit y (m d ) porosity (%) all wells thisted-3: low-permeability group thisted-3: high-permeability group regional trend (all wells) trend (thisted-3: low-perm. group) trend (thisted-3: high-perm. group) fig. 4. porosity–permeability plot based on conventional core analysis data from the gassum formation. 53 well-log interpretation and core description the evaluation of the reservoir quality of the gassum formation is primarily based on wireline logs and core-analysis data from thisted-1–4 (fig. 3). the quality of the logs is fairly good, but at certain intervals indications of caving lead to uncertain porosity estimates. lithology, clay content, porosity and permeability are evaluated from combined analyses of log data, lithological descriptions of cores and core analysis data. high, well test-based, porosities and permeabilities for the thisted wells were validated by gas permeability measurements from core analysis and are possibly caused by the relatively shallow depth of the gassum formation (hjuler et al. 2013). the regional porosity–permeability trend (fig. 4) significantly underestimates the actual (well test) permeability and was therefore calibrated to more closely approximate the well test results. evaluation of the sandstones indicates high average porosities exceeding 28% for the thisted-1 and 3 wells (table 1). the interpreted porosity of 25.3% for the thisted-2 well seems too low considering the fairly high test permeability of 3670 md; this is possibly due to erroneous log data. despite considerable variation of the assessed net sand thickness within the study area all thicknesses are considered sufficient for geothermal exploitation (table 1). detailed sedimentological descriptions and interpretations of the cores from the thisted-3 well indicate that the reservoir sandstones were mainly deposited in fluvial and estuarine environments (nielsen 2003; hjuler et al. 2013). considerable lateral and stratigraphical variations of the logs from the thisted wells indicate spatial and temporal variations of the sedimentary environment, which must have a significant influence on reservoir distribution and quality. it is thus important for the mapping of reservoir intervals to sb9 sb7 sb6 sb3 sb5 sb4 sequence boundary thisted-1 thisted-4 thisted-2 thisted-3 sandstone siltstone shale marlstone limestone dolomite coal gr dt gr caliper dt nphi rhob prs perm_log phie gr caliper dt nphi rhob prs perm_log phie core 2 core 3 core 4 core 5 datum line 1 g as su m f or m at io n fluvial estuarine shoreface offshore 1 4 2 3 5 g as su m f or m at io n g as su m f or m at io n g as su m f or m at io n depth (m) 750 800 850 900 depth (m) 1150 1200 1250 1300 depth (m) 1100 1150 1200 depth (m) 750 800 850 gr caliper dt rhob prs perm_log phie fig. 5. log correlation of sedimentary units within the gassum formation in the four thisted wells. 1–5: sequences, red lines: sequence boundaries (corresponding to incised valley floors), prs: potential reservoir sand. the sediments of the incised valleys are fluvial and estuarine. thisted-1 137 94.4 84.4 0.62 28.3 2500 3125 thisted-2 135 94.9 70.3 0.52 25.3 1300 1625 thisted-3 115 67.4 54.5 0.47 28.6 3500 4375 thisted-4 114 36.6 – – – – – table 1. reservoir parameters for net sand from the gassum formation* well formation gross sand net sand n/g† average estimated estimated no. thickness (m) thickness (m) thickness (m) porosity (%) gas perm. (md) reservoir perm. (md)‡ * shale and porosity cut-off applied, net sand is defined as sandstone with <30% shale and porosity >15%. † n/g: net sand thickness divided by formation thickness. ‡ estimated reservoir permeability: estimated gas permeability multiplied by 1.25 (upscaling factor). 5454 understand the variations in the sedimentary environment. hence the logs of the four wells have been correlated (fig. 5). five periods of sea-level rises and falls were identified within the gassum formation. during sea-level falls, incised valleys formed which were filled with fluvial and estuarine sediments during subsequent rises in sea level. the extent of reservoir sandstone deposits is limited by the extent of the valleys. to estimate the dimensions of sandstone bodies in the thisted area the extent of present-day estuaries has been studied. it was found that the width of estuarine sediment bodies exceeds 8 km and the lengths are probably several tens of kilometres. in the late triassic – early jurassic the overall shoreline was nw–se oriented. the seismic data do not allow the determination of the relief and extent of the incised valleys, but an overall ne–sw-orientation of the estuarine sediment bodies is assumed. the reservoir lithology is limited to fluvial and estuarine sandstones with the fluvial sandstones being c. 20 times more permeable than the estuarine sandstones because detrital clay clogs the pores of the latter. as expected, the reservoir properties of clayand siltstones outside the estuaries are poor and thus fluvial sandstones in the valleys may be hydraulically separated from each other. log correlation implies that the estuarine deposits may coalesce into larger reservoir units, or may be partly separated by layers of low permeability. in a more landward position towards the ne fluvial reservoir sandstones may amalgamate into more coherent reservoir units, leading to improved reservoir continuity. a reservoir model petrel® software was used to establish a reservoir model that covers an area of 11 × 13.5 km; flow simulations and dynamic modelling were performed using the eclipse® 100 software (fig. 6). input data for the reservoir model were prepared by integration of interpreted seismic data, depth structure maps, temperature assessment, petrophysical evaluation, core description and log correlations as described above. comparison with historic data was not possible as no pressure data of the thisted-2 and 3 wells are available; the simulated response from the dynamic modelling could only be adjusted because no significant temperature drop has been recorded in the production water since 1984. hence the model results are uncertain and should be evaluated with caution due to lack of calibration data. with this limitation in mind, and using the current well configuration, the model simulations indicate another 15 years of production with no alarming decrease in water temperature (less than 1°c). with thisted-5a as a new production well situated downflank of the existing well pair, the production could benefit from a higher temperature with no sign of cooler water arriving during an operational time of 30 years (uncertainty in modelling results as above). with thisted-5b or -5c as a new injection well situated at shallower depth than the existing well pair, and with thisted-2 still used as production well, the new injection well can benefit from higher permeability. concluding remarks by integrating data from geological, geophysical and petrophysical disciplines, a detailed geological model was established that provides a structural framework and demonstrates how depositional environment controls the lateral extent, continuity and quality of reservoirs. a detailed reservoir model was also established, which can simulate the remaining lifetime of the geothermal plant with the current well configuration, or the lifespan of the plant with the addition of a new production or injection well. the combination of geological and reservoir models provides indications of the best sites for additional production or injection wells. references hjuler, m.l., vosgerau, h., nielsen, c.m., frykman, p., kristensen, l., mathiesen, a., bidstrup, t. & nielsen, l.h. 2013: assessment of potential capacity increase of the geothermal plant at thisted by adding a new geothermal well. danmarks og grønlands geologiske undersøgelse rapport 2013/80, 90 pp. nielsen, l.h. 2003: late triassic – jurassic development of the danish basin and fennoscandian border zone, southern scandinavia. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 459–526. vejbæk, o.v. & britze, p. 1994: geological map of denmark 1:750 000. top pre-zechstein (two-way traveltime and depth). danmarks geologiske undersøgelse kortserie 45, 8 pp. fig. 6. permeability model that crosses the thisted-2 and 3 wells. thisted-2 thisted-3 top gassum fm base gassum fme 1000 100 10 1 permeability (md) 1 km w authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark; e-mail: mlh@geus.dk untitled 60 structural description of sections the rubjerg knude glaciotectonic complex is differentiated into 13 structural sections, which are named after the localities recorded in the geological crosssection of lønstrup klint by jessen (1918). the sections can be grouped into three zones within the complex: a distal zone (three sections farthest to the south), a central zone (seven sections in the middle part of the complex), and a proximal zone (three sections farthest to the north). they are named, from south (near nørre lyngby) to north (at lønstrup): ulstrup, stensnæs, martørv bakker, kramrende, brede rende, sandrende, stenstue rende, grønne rende, rubjerg knude fyr, stortorn, moserende, mårup kirke and ribjerg sections. three criteria were used for defining the sections: (1) the sections had to be bordered by marked footwall ramps, (2) each section should be characterised by uniform structural architecture, and (3) the sections had to be descriptively and geographically delimited. a good example of the first criteria is the steep thrust fault separating the sandrende section from the stenstue rende section. as an example of a uniform architecture, the grønne rende section can be mentioned, and finally the mårup kirke section includes the long barren stretch from the mårup church to ribjerg at lønstrup where the lack of geographical markers as well as characteristic footwall ramps is significant. each section is described separately, a general physiographic introduction being followed by four parts: (1) tectonic architecture, (2) sedimentary units, (3) structures, and (4) interpretation of structural development. the first descriptive part provides a general description of the macro-structures. the second part presents the sedimentary deposits, and although to some degree it repeats the lithostratigraphic descriptions of the formations (see above), the detailed observations of syntectonic sedimentation are relevant to an appreciation of the structural development in the individual sections. the third part concerns the description of mesoand mini-structures (thrust faults, folds, faults, shear zones, joints, fractures, breccias, polydiapirs etc.). the description of each section is concluded with an interpretation of the formation of the structures. this interpretation should not be confused with the overall interpretation of the dynamic development of the progressive thrust-fault deformation that follows the systematic descriptions of the sections. the organisation of the section descriptions follows the general systematics of structural geology: the description of geometry, the kinematic investigation, interpretation of the dynamics and finally the analysis can be concluded by a tectonic synthesis (dennis 1972). the reader should note that the structural elements are numbered from distal to proximal. this is a consequence of the systematic analysis; in order to obtain an overview, the reader can compare the structural description given for each section with the dynamic development presented in the latter part of this bulletin. ulstrup section the southern frontal edge of the rubjerg knude glaciotectonic complex is positioned in the ulstrup section where the undeformed foreland is exposed below the hanging-wall flat of the last displaced and most distal thrust sheet (ul01, see plate 2). one of the most interesting problems addressed in the analysis of this complex is the presence of two long thin thrust sheets that were translated southwards for about 500 m from their ramps to their present positions without complete internal disruption. the ulstrup section represents the foreland margin of the thrust-fault complex, and the two flat-lying thrust sheets extend from the southern edge of the thrust front at tvonnet rende to the steep ramp thrust separating the ulstrup and stensnæs sections (plate 1). the stensnæs ramp thrust was initially regarded as the foreland thrust (pedersen 1987), but new outcrops of the southern ulstrup thrust exposed in 1996 and 1997 clearly demonstrated additional details of the frontal thrusting. consequently, the long cliff section showing horizontal bedding that had previously been regarded as a primary, undeformed sedimentary unit (fig. 42) is now interpreted as a displaced thrust sheet (ul01). the ulstrup section is truncated by a broadly horizontal glaciotectonic and erosional unconformity above which aeolian sand was deposited, either as sheet sands or as small dunes, up to 10 m high. 61 tectonic architecture the ulstrup section comprises the two flat-lying thrust sheets, ul01 and ul02. the tip line of the hangingwall ramp (the edge of the frontal thrust) is situated on the northern side of tvonnet rende (for location, see plate 1). unfortunately, the precise position is obscured by late, syntectonic erosion and sedimentation at the front of the thrust-fault complex, as well as by sand scree covering the outcrops at tvonnet rende. the northern boundary of the section is the footwall ramp and flat of the ul02 thrust sheet, which forms a transitional zone of imbricate thrusting related to the frontal part of the stensnæs section. the frontal part of the ul01 thrust sheet was displaced from the upper footwall ramp (fig. 43) along an upper footwall flat on the top surface of the foreland. at the tip of the thrust sheet, the thrust fault dips gently towards the foreland, and a small foreland-dipping ramp is also located at point 6040 m in the crosssection (plate 1, see fig. 50). the ramps probably formed due to erosion in front of the propagating thrust-fault edge. in the central part of the ul01 thrust sheet, a synform structure is associated with a chaotic breccia, interpreted as the collapse of a frost mound or sandmud diapir below the thrust fault; the synformal depression is referred to hereafter as ulstrup rende (for location, see plate 1). the trailing end of the thrust sheet starts at the upper footwall ramp of the foreland from where the hanging-wall flat is inferred to continue along the footwall flat to the footwall ramp at its trailing end. the total length of the thrust sheet is about 750 m and the displacement is estimated at 350 m, which is the distance from the footwall ramp at point 5800 m in the cross-section to the frontal termination in tvonnet rende (plate 1). the thickness of the thrust sheet varies from 10 to 20 m, decreasing towards the tip to the south, and increasing in thickness where syntectonic deposits fill the piggyback basin on the back of the thrust sheet. the thickness of the lønstrup klint formation is only 6 m in the southern part of the piggyback basin due to erosion related to elevation during thrust faulting. the ul02 thrust sheet is 600 m long, with the frontal edge situated on the northern side of the ul01 piggyback basin at point 5900 m, and the trailing end disappearing into the décollement zone at point 5300 m. the thrust fault consists of a more than 400 m long footwall flat on top of ul01 extending from the footwall ramp hinge at point 5360 m southwards to the fig. 42. the steep sandy cliff of the ulstrup section displays horizontal bedding of an apparently undisturbed deposit. however, structural analysis of thrust-fault relationships to the south indicates that it is a long flat thrust sheet displaced more than 500 m towards the foreland to the south. the cliff section is 25 m high and the view is towards the south. photograph: june 1984. 62 gently dipping frontal bend at point 5780 m in the cross-section (plate 1). the frontal bend corresponds to the hinterland-dipping limb of the flat-topped hanging-wall anticline of ul01. this hanging-wall anticline compares well with the structure of the thrust-fault model in fig. 6, and the southernmost c. 100 m of ul02 can be regarded to have been emplaced piggyback on ul01 during the translation over the frontal footwall ramp. the ul02 thrust sheet is only 5–6 m thick above the flat-topped anticline, whereas the thickness increases to 20 m at the trailing-end ramp. the displacement along the northern thrust fault is 550 m, with an uncertainty of 10–25 m depending on the interpretation of the shape of the trailing-end ramp and the amount of erosion of the frontal part at the piggyback basin at ulstrup rende. sedimentary units the sedimentary units in the ulstrup section comprise the upper part of the lønstrup klint formation, erosional remnants of the lower part of the rubjerg knude formation, and a variety of intercalations of the rubjerg knude formation distinguished here as the ulstrup beds. the l/r-unconformity between the lønstrup klint and rubjerg knude formations can be traced along the upper part of the ul01 thrust sheet in which it also forms the base of the piggyback basin. a few younger erosional unconformities, below and above the ulstrup beds, are of only local significance within the ulstrup section. lønstrup klint formation the mud-dominated lønstrup klint formation forms the main part of the thrust sheets in the ulstrup section, and has an average thickness of about 10 m (fig. 19). in the southern thrust sheet, the fine-grained sand beds are thin and only small-scale current ripples have been observed. a combination of load structures (balland-pillow) and water-escape structures (convolution and small-scale diapirs) are developed at certain horizons. above the hanging-wall flat, a zone about 2 m thick takes the form of a mobilised mud breccia, which can be characterised as a sole thrust zone. this breccia is superimposed by beds affected by a brittle type of brecciation, forming cracks and joints in an up to 4 m thick zone in the lower part of the thrust sheet. fig. 43. the upper footwall ramp of the foreland (foreland fwr) along which the ul01 thrust fault propagated (ul01hwf: ulstrup thrust sheet 01 hanging-wall flat), and from where it continued for more than 300 m over the footwall flat of the foreland. the clif f section is 25 m high and south is to the right. photograph: may 1998. 63 the northern thrust sheet (ul02) displays a more sandy part of the lønstrup klint formation (fig. 25). here the sand beds are up to 1 m thick and waterescape structures, convolute bedding and flame structures commonly disturb the primary bedding. balland-pillow structures are more common towards the trailing end of the thrust sheets. rubjerg knude formation in the ulstrup section, the rubjerg knude formation comprises three different depositional units: the main ‘background’ sedimentation of outwash sand, the glaciolacustrine ulstrup beds (fig. 25), and the coarsegrained glaciofluvial ulstrup beds (fig. 20). the main depositional unit is fineto mediumgrained meltwater sand represented in the footwall block of the foreland (fig. 43). small-scale current ripples occur in the parallel bedded sand, which is interlayered with c. 0.5 m thick trough cross-stratified beds. the glaciolacustrine ulstrup beds form a 3–5 m thick unit that is only found on the back of the ul02 thrust sheet (fig. 25). this unit consists of dark bluish grey, laminated clayey mud interlayered with a few sandy beds up to 0.5 m in thickness. the unit was deposited on a bedding-parallel unconformity, which is only discordantly developed in the northernmost trailing part of thrust sheet ul02. the unit thins out towards the south, and disappears near the hinterland-dipping limb of the flat-topped hanging-wall anticline formed above the footwall ramp of the foreland. intraformafig. 44. the mud-mobilised thrust-zone breccia consists of structureless mud with scattered clasts floating in the matrix. dilation cracks filled with sand are superimposed on the mud-mobilised brecciation fabric. this reflects two phase of cataclastic deformation: a first phase of water-over-pressured brecciation (hydrodynamic brecciation), and a second phase of brittle fracturing when the mud was consolidated, dehydrated or perhaps frozen. handle of spade is 12 cm. photograph: june 1997. tional hydrodynamic brecciation, including small-scale diapirs and slump-like features, deformed the clayey mud; such deformation is not seen in the lønstrup klint formation below the unit. the glaciofluvial ulstrup beds occur in the ulstrup rende depression between points 5900 and 6000 m in the cross-section (plate 1). this unit is a c. 8 m thick succession of meltwater gravel fining up into coarse-grained sand (fig. 20). large-scale trough crossbedding dominates the succession and clasts up to boulder size occur in the lowermost 2 m (fig. 28). lithologically, the clasts are dominated by flint, but clasts of fossil frozen sand are abundant (fig. 29). the glaciofluvial ulstrup beds are overlain by c. 5 m of mediumto coarse-grained sand of the rubjerg knude formation. on the north side of the ulstrup rende depression, between points 5900 and 6000 m, slump-folded sand beds and sedimentary breccias occur in the upper part of this succession, suggesting gravity gliding down the steep slope of the depression (piggyback basin of ul01). structures and breccias the most important structures related to the thrust faults in the ulstrup section are the breccias occurring in the thrust zones above the thrust-fault surfaces. they appear to have formed by collapse of the thrust sheet during translation. the low-angle anastomosing faults that developed in the most distal part of the thrustfault complex originated similarly during translation 64 and are associated with southerly dipping normal faults. a significant collapse structure that formed beneath the ulstrup rende depression is also worthy of note. thrust-zone breccias the thrust-zone breccias occur above the hangingwall flat of the ul01 and ul02 thrust sheets, where they affect the mud-dominated lønstrup klint formation. the thrust zone is up to 4 m thick in the most distal part of the thrust-fault system (southern part of ul01), and decreases in thickness to 1 m northwards; it can be traced along the hanging-wall flat of ul02 for a considerable distance. the thrust-zone breccia consists of mobilised mud with irregular clasts of mud scattered throughout (fig. 44). some patches may be more sandy and others more clayey, and lenses and layers of sand may be present. the mobilisation was apparently initiated as sandy mud-fluid that developed at the thrust-fault surface and extended up into the sedimentary unit (fig. 45). in many cases, the initial hydrodynamic brecciation of the thrust zone left segments along the displacement surface of the thrust fault, which were modified and developed into elongated cataclasts along the sole of the breccia zone. convolute bedding and small-scale diapirism are also present. the mobilised mud was subsequently transected by dilation cracks and sand-filled fissures (fig. 46). the dilation cracks form an irregular network and the sand-fill was injected into consolidated mud (figs 44, 46, 47). more or less horizontal sand-filled cracks have been observed in the frontal part of the ul01 thrust sheet, where they are up to 15 cm thick and appear up to 1 m above the hanging-wall flat. the sand in the cracks shows planar horizontal lamination and small-scale current ripples and a few vertical sandfilled pipes extend upwards from the cracks (fig. 48). towards the frontal tip of the ul01 thrust sheet, an increasing number of bedding parallel or low-angle anastomosing fractures and small-scale faults appear to be related to an increased rate of internal gliding. this is an indication of how close the thrust sheet was to disintegration and a loss of cohesion. zones 0.2–0.5 m thick, grading into mobilised mud, occur in between the anastomosing fractures, resulting in the destruction of bedding (fig. 49). foreland-dipping hanging-wall flat faults in the frontal part of the ul01 thrust zone, forelanddipping faults become increasingly common. these faults are either foreland-dipping (20–30°s) ramps formed by the hanging-wall flat scouring into the footwall flat (fig. 50), or sets of 50°s dipping normal faults with displacements of about 10 cm. these structures are considered to be the result of partial collapse of the tip of the foreland-dipping limb of the hangingwall ramp above a low-angle hanging-wall ramp translated along the footwall flat of the foreland. intrusive contact thrust-fault surface fig. 45. mud mobilisation along the hanging-wall flat of the ul02 thrust sheet. from the thrust-fault sur face, a sandy mud fluid intruded along fractures and up into the lower part of the thrust sheet where it formed a mud-intrusion. during the initial hydrodynamic brecciation of the thrust zone, small relict segments remained at the thrust plane where they were modified and developed into elongate cataclasts in the sole of the thrust breccia zone. photograph: july 1998; matchbox for scale. 65 fig. 46. subsequent to mud mobilisation, the thrust zone was transected by dilation cracks and fissures, which form an irregular network into which sand was ‘injected’. the mobilised mud had clearly become consolidated before the sand-filled cracks formed. the close-up inset illustrates the ‘reverse’ drag of the sand-fill (arrowed) indicating an upward direction of flow in the fissure. photograph: may 1998. 66 collapse structure in the ulstrup rende structures in the central part of ulstrup rende (figs 51, 52) are interpreted to represent a collapsed diapir. the early phase structures include thinning of the lønstrup klint formation in the thrust sheet, and formation of concave troughs in its surface. vertically or steeply dipping sand breccias with upward directed flow structures cross-cut the thrust-zone breccia. the appearance of structureless sand pockets indicates sandfill of mobilised sediment from an over-pressured zone in the subsurface. the complex of breccias and reorientated bedding is interpreted as a collapse structure; it is considered to be responsible for the formation of the ulstrup rende depression, and the disruption of the thrust sheet along steeply dipping fractures. in the breccia zone, steeply dipping sand-filled cracks and normal faults formed prior to the continued deposition of the rubjerg knude formation in the depression. on the southern side of the ulstrup rende depression, the muddy part of the thrust sheet is displaced by steeply dipping normal faults. downthrow is to the south, synthetic towards the depression, and the faults are thought to be related to the collapse of the diapir structure. interpretation of structural development there are two reasons why the ulstrup section deserves special attention. the first is that it represents a foreland thrust section with long lateral transport of thin thrust sheets, the nature of which has not previously been documented. secondly, it demonstrates the likely development of the initial stages of deformation, which the remainder of the thin-skinned thrust faulting in the rubjerg knude glaciotectonic complex also experienced before the uppermost part was eroded. thus the first phase of thrust-fault deformation is preserved here whereas it is almost never represented in the thrust-fault sections that have been fig. 47. the formation of thrust-zone breccias in the distal part of the thrust-fault complex is here illustrated in four stages of development. (1) the initial undeformed sediment (lønstrup klint formation) comprises clayey mud interlayered with thin sand beds. (2) a mud-fluid is formed above the thrust-fault surface (line with open triangles ) from where it is injected up into the layers above (see fig. 45). (3) increasing mud-mobilisation results in the formation of a structureless matrix with dispersed matrixsupported clasts of the primary sediment (see fig. 44). note the small normal faults indicating an on-going process of collapse. (4) the mobilised mud becomes consolidated and the thrust-zone breccia develops into a more brittle stage; dilation cracks form into which water-saturated sand is injected (see fig. 46). 67 more intensely deformed. the interpretation of the structural development can be summarised in the following nine stages. 1. initial thrust-fault fracturing and thrust-fault propagation took place during mobilisation of mud along the hanging-wall flat. at this stage, the thrust-zone breccia was formed due to high pore-water pressure in an unfrozen stage. 2. the ul01 and ul02 thrust sheets probably started to move along the décollement zone as one coherent thrust sheet, and first separated into two thrust sheets after the frontal part of the sheet had passed the most distal foreland footwall ramp. 3. the ramping up of the northern ul02 thrust sheet probably increased the pore-water pressure, and when this increase also affected the frontal part of ul01, diapirism was initiated under the central part of the southern thrust sheet. 4. the diapiric uplift and erosion took place in the elevated surface. this erosion extended through the rubjerg knude formation to locally intersect the l/r-unconformity. the residual coarse clastic gravel fig. 48. a subhorizontal sand-filled fracture occurring in the frontal part of the ul01 thrust sheet. the sand-filled fracture appears about 1 m above the hanging-wall flat; the sand shows planar horizontal lamination and current ripple cross-lamination. the sand-filled fracture is interpreted to have formed during ground-frozen conditions, whereas the vertical sand-filled pipe probably reflects loading of the thrust sheet when it ultimately lost its carrying pore-water pressure and settled on its hanging-wall flat. photograph: june 1997. fig. 49. subhorizontal anastomosing faulting (centre left) with mud-mobilisation developed in domains between fault fractures in the frontal part of the ul01 thrust sheet. photograph: june 1997. 68 fig. 50. a: foreland-dipping ramping of hanging-wall flat formed by scouring-erosion of the footwall flat into the top surface of the foreland. note in the close-up (b) that some hydrodynamic brecciation occurred in the footwall ramp just below the thrust zone. photograph: june 1997. on the footwall flat at the surface of the foreland in the ulstrup section. 7. the consequence of the diapir collapse was the formation of the depression in ulstrup rende. redeposited coarse-grained clastic material filled the depression, generating the glaciofluvial ulstrup beds. the occurrence of fossil frozen-sand clasts implies that part of the surface, the rubjerg knude formation, was ground-frozen – probably that part of the thrust sheet that had been elevated due to the propagation up over the central ramp. the ground-frozen condition was probably also responsible for freezing of the mobilised mud in the thrust zone and the subsequent development of sand-filled dilation cracks (fig. 47) bed on the unconformity surface was probably the original source for the large amount of coarse-grained material in the ulstrup rende depression. 5. the mud-sand volcano broke through to the surface, and the mobilised mud and sand were extruded with the release of the high pore-water pressure. 6. the surface of the diapir collapsed and gravel and sand filled the fractured structure. the collapse was probably contemporaneous with the loss of pore pressure throughout the thrust zone. the release of over-pressure in the thrust zone resulted in the final settling of the thrust sheet 69 8. the depression was ultimately filled by the sand of the upper unit of the rubjerg knude formation. the frontal edge of the northern thrust sheet propagated towards the northern side of ulstrup rende and parts of its leading tip slumped down the steep slope of the depression. this indicates the sequential and later movement of the northern thrust sheet. the depression can be interpreted in part as a piggyback basin that formed according to the model demonstrated in fig. 7. 9. finally, the uppermost sediments of the rubjerg knude formation covered the section before the glacier advanced across the area. the uppermost metre of sand was transformed into a glaciofluvialsand-glacitectonite. fig. 51. the ulstrup rende depression is interpreted to represent a combination of a piggyback basin (according to the model in fig. 7) and the collapse of an underlying mud diapir or frost-mound feature. note the steep normal fault on the left side of the depression (to the north). photograph: june 1997; spade (lower centre) for scale. fig. 52. chaotic sand/gravel breccia in the centre of the ulstrup rende depression, which is interpreted as the result of the collapse of the diapiric structure created in the subsurface below the ul01 thrust sheet. photograph: may 1998. 70 stensnæs section the stensnæs section is named after stensnæs, which is a minor point (danish: sten = stone; næs (pynt) = point) at a gentle bend in the cliff section. stones and erratic blocks, probably derived from the up to 2 m thick sandy till and the erosional unconformity below the vendsyssel formation, were formerly abundant on this part of the beach. the sandy till thins out southwards where a glaciotectonic unconformity truncates the section. the stensnæs section displays the most spectacular folds in the rubjerg knude glaciotectonic complex (fig. 53). the folds are situated at the transition from the flat-lying beds to the south (the ulstrup section) and the main thrust-fault imbrications to the north. the fold complex is truncated by an erosional unconformity forming a depression in which slump slides and sedimentary breccias derived from the tip of the thrust sheet were deposited after collapse and gravity gliding. in general, the fold complex is well exposed, whereas the transition further southwards is often covered by sand scree. during the years of study of the rubjerg knude glaciotectonic complex by the author, variations in the degree of exposure of the stensnæs section have contributed to a fuller understanding of the structural development of the section that represents the foreland margin of the thrust-fault complex. tectonic architecture the stensnæs section comprises four thrust sheets (sn01, sn02, sn03 and sn04, plate 2). the southern boundary of the section is defined by the footwall ramp of ul02, although the trailing end of ul02 is here included in the description of the imbricate duplex that hosts the fold complex. the northern boundary is the hanging-wall ramp of mb01/mb02, that is thrust up along the footwall block of sn04. the thrust sheets comprise the uppermost part of the lønstrup klint formation and a relatively thin cover of the overlying rubjerg knude formation sediments. two discrete piggyback basins (early and late) were formed above the sn01 and sn02 thrust sheets. the sn01 thrust sheet is about 30 m thick, and although the thrust sheet includes a number of small duplex imbricates it can be subdivided into upper and lower segments. the lower segment is thrust onto the ul02 footwall block with a displacement of about 30 m. the displacement of the upper segment is partitioned fig. 53. the fold complex developed in the transitional imbricate zone between ulstrup and stensnæs sections. note that the thrust faults acting as flexural slip surfaces in the folding continue into the bedding-parallel thrust-fault flats to the south (right). photograph: july 2000. 71 into a series of minor dif ferential displacements ranging from minor flexural slips along bedding surfaces in the fold structures to imbricate offsets of about 1–3 m. in the frontal part of the sn02 thrust sheet, the lønstrup klint formation is only a few metres thick, and is overlain by 10 m of sediments of the rubjerg knude formation resting on the l/r-unconformity. at the trailing end of the thrust sheet, the lønstrup klint formation is more than 15 m thick, whereas the rubjerg knude formation is cut off by the footwall ramp of sn02 (corresponding to the hanging-wall thrust fault for the imbricates of sn03). the accumulated displacement of sn02 amounts to 88 m, including 15 m up over the footwall ramp and 73 m along the footwall flat of sn01. the sn03 thrust sheet may be divided into four imbricate segments thrust onto the footwall ramp of sn02. the displacement along the hanging-wall thrust fault is 45 m, but it was also carried piggyback on sn02 during the translation along the décollement zone, which gives an accumulated displacement of 163 m for the sn03 thrust sheet. fig. 54. ball-and-pillow structures in the lønstrup klint formation in the stensnæs section. note that the size of the structures reflects the thickness of the sand beds involved. photograph: july 2000; staff divisions are 20 cm. fig. 55. dish-and-pillar structures developed in the lowermost sandy bed above the hanging-wall flat of the sn2 thrust sheet in the stensnæs section. this type of water-escape structure is interpreted to have formed during the thrust-fault translation along a hangingwall flat due to the high water pressure released from the sole of the thrust sheet. photograph: july 1998. 72 the sn04 thrust sheet is about 20 m thick, including an up to 8 m thick unit of the rubjerg knude formation resting on the l/r-unconformity above the lønstrup klint formation; this unit increases in thickness to about 15 m towards the trailing end of sn04. total displacement is about 45 m, including a hanging-wall ramp-and-flat propagation along the footwall ramp of sn03. sedimentary units the sedimentary units that crop out in the stensnæs section are dominated by the upper sandy parts of the lønstrup klint formation. a lens of a glaciolacustrine diamictite is preserved as an imbricate thrust sheet, c. 3.5 m thick and 22 m long, and is tentatively interpreted to form part of the rubjerg knude formation. the latter is mainly represented by a 6–8 m thick succession including the residual gravel deposited on the uneven surface of the l/r-unconformity. an upper unconformity truncates the sn01 and sn02 thrust sheets, and coarse gravel beds as well as olistoliths of the lønstrup klint formation were deposited in a late piggyback basin (snstru in plate 2). lønstrup klint formation the sedimentology of the lønstrup klint formation in the stensnæs section is similar to that described for the ulstrup section. the maximum thickness of the formation is about 15 m, and lithologies are dominated by 0.2–0.8 m thick fine-grained sand beds with climbing ripple lamination. thin laminae of mud and small amounts of detrital organic material commonly drape the ripples. the sand beds are interlayered with thin laminated dark grey mud beds; micro-faulting is evident in this facies. ball-and-pillow and convolute structures are abundant (fig. 54). in the uppermost 4 m of the formation, the sandy beds become thinner and the uppermost part is dominated by mud. in sand beds inferred to be situated above the hanging-wall flat, dish-and-pillar structures have been observed (fig. 55) that are interpreted to have formed by water-escape processes related to the thrusting. cheel & rust (1986) provided a model for the development of water-escape structures in glacial outwash deposits from ottawa, canada. in their model, a sequential development from simple load structures through detached ball-and-pillow structures to dish structures is demonstrated. the model predicts a strafig. 56. the piggyback basin situated above the erosional truncation of thrust sheet ul02 and sn01 in the stensnæs section (snstru in plate 2). note that normal listric faults (arrow) may be traced to the trailing end of the large olistoliths (green) deposited in the meltwater sand and gravel. this is interpreted as the result of gravity gliding of the tip of a thrust sheet, which during propagation from the north collapsed subsequent to thrust faulting up above the mean level of sedimentation. photograph: july 2000; thrusts indicated in red, unconformity indicated in purple . 73 tified distribution with convolute stratification at the base of a bed, ball-and-pillow structures dominating the main part, and dish structures formed in the uppermost part of the bed (or unit) resulting from excess pore-water fluid pressure. the observed waterescape structures in the stensnæs section as well as in the sections further to the north compare well with the model of cheel & rust (1986) (figs 54, 55; see also fig. 77), although they suggested the triggering mechanism to be earthquakes or movements due to melting of dead-ice. rubjerg knude formation in the stensnæs section, two units of the rubjerg knude formation are distinguished: (1) a lower unit dominated by fineto medium-grained glaciofluvial sand, and (2) an upper unit of varied sedimentary breccias and coarse-grained clastic deposits that infills the piggyback basin on top of the truncated thrust sheets in the section (fig. 56). the sand deposited in the lower unit may be planar parallel stratified, or exhibit large-scale trough cross-bedding with shallow troughs, only 0.5 m deep. the lower unit of the rubjerg knude formation is estimated to be 6–8 m thick, and although strongly affected by hydrodynamic brecciation, the unit compares well with the description of the lower part of the formation provided by sadolin et al. (1997). the upper unit was deposited on the erosional unconformity truncating the back of the ul02 thrust sheet, the fold and imbricate complex of sn01, and the tip of the sn02 thrust sheet (snstru in plate 2). the central part of the basin is about 7 m thick, decreasing towards both the north and south forming a relatively narrow trough. the sediments in the basin consist of large-scale irregularly trough cross-bedded glaciofluvial sand and gravel. blocks of sandy mud, which can be identified as derived from the lønstrup klint formation, were deposited as sedimentary breccias in the basin. the clasts are up to 1 × 5 m in size; blocks and clasts less than one metre in size were commonly rotated during redeposition. normal listric faults in the northern part of the basin relate to the deposition of the largest olistoliths (fig. 56). the southern part of the basin was tilted subsequent to deposition due to the fault-bend folding of the hanging-wall flat of ul02 when it propagated along the footwall flat during the latest phase of thrust faulting. thus the upper unit of the rubjerg knude formation in the stensnæs section is interpreted to have been deposited in a piggyback basin in which thrust-sheet tips thrust up from the north collapsed and gravity-glided out into the basin. structures two types of structures in the stensnæs section are related to thrust faulting: (1) the folding related to duplex imbricates, and (2) extensional faults related to push-from-the-rear in the trailing end of a thrust sheet. imbricate duplex folding the folds in the stensnæs section can be described as flexural slip folds, and have amplitudes of 1–3 m and wavelengths of 2–5 m (figs 53, 57, 58). the folds are very irregular in shape, however, and cannot be explained in terms of simple compression. analysis of the flexural slip surfaces shows that the folded layers were separated into segments and that discordant relationships exist between beds in neighbouring segments. by defining the segments as small imbricate thrust sheets in a duplex, thrust-fault terminology can be applied and hanging-wall and footwall thrusts of the individual duplex segments defined. in fig. 57 this has been done by identifying the footwall ramps (fwr), thus distinguishing five imbricate thrust sheets about 1 m in thickness. the folds in the imbricate duplex segments include both hanging-wall anticlines and footwall synclines. as documented by the refolding of the upper hanging-wall anticline in fig. 57, the folds are superimposed by sequential phases of folding and thus also phases of imbricate thrusting. figures 58 and 59 illustrate the sequential development of imbricate thrusting; the imbrication steps forward towards the footwall ramp of ul02 to the south. note also that the hanging-wall flat of each imbricate continues into an intraformational bedding-parallel thrust fault. this can be difficult to recognise in an isolated exposure, where the flats cannot be traced back to the ramp structures in the imbricate duplex (fig. 53). extensional faults in the trailing part of the sn03 thrust sheet, listric extensional faults have been observed (fig. 60). displacements along the faults are up to 0.5 m and the 74 fig. 57. flexural slip folding in the sandy beds of the lønstrup klint formation in the stensnæs section. numbers (1–4) refer to the sequential development of the thrust faults in the duplex structure, which probably formed during the collapse of the footwall ramp to the south (ul02fwr). note that the ramp thrust faults propagate into bedding-parallel flats. photograph: may 1996. fig. 58. detail of the fold structures formed by superimposed ramp propagation (see fig. 57). the thrust-fault segments in the imbricate duplex are marked by black lines and the recognised footwall ramps are annotated fwr. hanging-wall anticlines are dominant, but footwall synclines also add to the fold framework. photograph: may 1996. 75 fig. 59. simplified model of superimposed folding formed by sequential imbrication in a duplex complex. step 1 shows the undeformed beds with the ramps bordering the segments indicated. the numbers refer to the sequentially propagating hangingwall ramps. step 2 shows the first hanging-wall anticline to be formed by the progressive collapse of the trailing end of the footwall block (footwall ramp collapse). step 3 illustrates the formation of an antiformal stack during the progressive superimposed deformation. step 4 demonstrates the developed stage of superimposed fault-bend folding comparable to the structures illustrated in figs 57, 58. 76 fig. 60. low-angle listric extensional faults in the sn03 thrust sheet, in the central part of the stensnæs complex. a: an overview of the macroscopic structure, where the sn04 thrust sheet ramps up along the footwall flat of sn03 and pushes it laterally in the back. the resultant extensional fault imbricates form a boudinage-like network. b: detail of the listric extensional faults (arrows) interpreted to have been formed by push-from-the-rear. photograph: july 1996; staf f divisions are 20 cm. 77 fig. 61. illustration of the difference between the mild deformation af fecting the top of the footwall block and the strong deformation of the hanging-wall block. only a 20 cm thick zone below the shear-laminated thrust-fault surface was affected by low-angle extensional faulting grading down into a minor normal fault network. it is thus evident that the elevated water pressures supporting the thrust sheet were transmitted to the hanging-wall flat, where intense hydrodynamic brecciation took place. photograph: july 1998. 78 spacing between the faults is 0.2–0.6 m, which creates a boudinage-like network. the structures are interpreted to have been formed by push and loading of a thrust sheet ramping the formerly monoclinal fault-bend-folded thrust sheet, which responded to the gravity spreading by displacements along the extensional faults. similar mini-scale extensional normal faults are recognised in the footwall block below the hanging-wall flat of sn03 (fig. 61). it is remarkable that the thrustfault deformation only weakly affects the top of the footwall block, while the hanging-wall block is strongly affected by hydrodynamic brecciation. a zone only 20 cm thick below the shear-laminated thrust-fault surface is affected by low-angle extensional faulting, and grades down into a minor normal fault network (fig. 61). isoclinal folding has been observed in the narrow shear-laminated thrust-fault zone, adding to the impression of high strain along the thrust fault. however, it is clear that the elevated water pressures supporting the thrust sheet were located in the hanging-wall flat. interpretation of structural development the formation of the imbricate duplex fold complex in the stensnæs section is evidently related to ramp propagation. two footwall ramps are significant: the footwall ramp of ul01, which can be regarded as representing the footwall ramp of the foreland, and the footwall ramp of the trailing end of ul02. the ul01 footwall ramp acted as a stopping block for the forward push of the imbricate thrust sheets, and the propagation of this ramp was responsible for the general gentle northerly tilt of the structures. the ul02 footwall ramp was subjected to successive thrust-fault splay formation, and consequently imbrication and superimposed folding, which can be viewed as the collapse of the trailing end of the ul02 thrust sheet. the imbricate duplex fold complex of the stensnæs section can be readily compared to the model for connecting splay duplexes of the sevier thrust belt in the cordilleran fold belt (mitra & sussmann 1997). there is a close similarity with respect to the growth of the duplex by successive connecting splays of the thrust fault and the creation of folds by thrust-fault propagation. moreover, the analysis of the imbricate duplex fold complex implies that the imbrication started at the footwall ramp in the trailing end of the system and propagated towards the foreland. it can therefore be argued that the process was one of footwall ramp collapse. a simplified model for the growth of duplexes in a connecting splay duplex system is illustrated in fig. 59. with respect to the thrust-fault displacement, an active and a passive stage of translation need to be distinguished. the active translation is the amount of displacement of the thrust sheet arising from propagation along its hanging-wall thrust fault. the passive translation is the amount of transport arising from the displacement of the underlying thrust sheet which carries it piggyback fashion. during this latter translation, the underlying thrust sheet may propagate footwall ramps, which will fold the actively translating thrust sheet as well as the piggyback thrust sheets into hanging-wall anticlines. this type of deformation will create an antiformal stack. the stensnæs imbricate duplex fold complex may thus also be viewed as a mesoto macroscopic-scale antiformal stack (fig. 59). a roof thrust, which is 110 m long and about 8–12 m thick covers the sn01–sn02 duplex. the accumulated length of the three duplex segments is 175 m; c. 65 m is thus missing in the balance calculation of the southern half of the stensnæs section. it is most likely that part of the initial thrust sheet has been eroded away, but up to c. 45 m of it might have been incorporated in a foreland-dipping frontal thrust structure preserved in the chaotic imbricate fold complex. martørv bakker section the name martørv bakker is derived from the peat exposed in the coastal cliff (danish: mar = sea; tørv = peat; bakker = hills), which is covered by aeolian sand dunes above the northern end of the section. in the southern part of the section, the vendsyssel formation forms the top unit in the clif f. the vendsyssel formation was deposited on an erosional unconformity above the glaciotectonic complex. all the posttectonic deposits are prone to cliff erosion and the resultant scree partly obscures the structures in the martørv bakker section. in addition, the unconformity at the base of the planar-bedded vendsyssel formation is a focus for groundwater seep which also conceals details in the exposures. however, two important architectural elements have been recognised: (1) the common appearance of minor duplexes in the southern part of the section, and (2) the occurrence of southerly dipping normal faults in the northern part. the southern boundary of the section is the footwall ramp of sn04 in the trailing end of the stensnæs 79 section. the northern boundary is not defined by a simple reference point in the cross-section, but by the trailing end of the mb04 thrust sheet which is a combination of footwall ramp and footwall flat below the leading-edge thrust fault of the kramrende section. tectonic architecture the martørv bakker section is subdivided into four thrust sheets (mb01–mb04). the thrust sheets in the southern part of the section are subdivided into upper and lower duplex segments of which only the lower duplex segments are distinguished by separate annotations (mb02u1–mb02u3, plate 2). in the frontal part of the section, smaller imbricate duplexes are associated with syntectonically formed hydrodynamic breccias and ball-and-pillow structures. the mb01 thrust sheet is up to 20 m thick and comprises the upper part of the lønstrup klint formation. the frontal hanging-wall ramp was thrust up along the 15° dipping footwall ramp of sn04, and displacement is estimated at about 55 m. the bedding becomes steeper in the trailing end of the thrust sheet, probably due to the relatively steep dip of the ramp in the subsurface from the 20 m to the 10 m flat level. the mb02 thrust sheet is long and flat-lying and occupies more than 400 m of the section. the frontal part is only 10–15 m thick and was displaced along its upper hanging-wall flat along the footwall flat of mb01 for a distance of about 180 m. the trailing part of the mb02 thrust sheet is 20 m thick, but as indicated in the balanced cross-section it roots down to the 30 m décollement level (plate 2). the thrust sheet is displaced by a prominent normal fault in the central part of the martørv bakker section (fig. 62). above the lønstrup klint formation, a marked basin developed in the hanging-wall block of the normal fault. the fig. 62. the normal fault developed in the central part of the martørv bakker section. in the hanging-wall block to the south, the fluvial sands of the lower part of the rubjerg knude formation are preserved in a ‘fault trap’ along the fault plane. diamict sediments were deposited above the sand in a piggyback basin that developed during the thrust faulting of the mb02 thrust sheet. note that the thickness of the lønstrup klint formation in the footwall block decreases downwards along the normal fault drag. this is interpreted as a foreland-dipping limb related to a hanging-wall anticline formed prior to offset by the normal faulting. the formation of the normal fault is interpreted to be related to a foreland-dipping limb of a hanging-wall anticline at the tip of a subsurface thrust-sheet segment (mb02u3 in plate 2). photograph: june 1993. 80 sediments in this basin were described as moraine sand by jessen (1918, 1931). the mb03 thrust sheet is 30 m thick and made up of the lønstrup klint formation, which is here strongly deformed by hydrodynamic brecciation and dislocated by a number of bedding-parallel minor thrust faults. the frontal part is flat-lying, whereas the dip of the bedding increases to 25–30° at the trailing end indicating a bend over an upper hinge on top of the footwall ramp of mb02. in the exposed part of the section, the mb03 thrust sheet is only about 120 m long, and it is inferred that the foreland-dipping structures in the frontal part reflect re-orientation due to hanging-wall ramp propagation along the footwall flat of mb02. the mb04 thrust sheet forms a flat-topped hanging-wall anticline above the footwall ramp of mb03. the thrust sheet is displaced along a normal fault parallel to the foreland-dipping bend in the top of the frontal part of mb03. the normal fault does not displace the footwall flat of mb02, and is therefore regarded as a structure related only to the framework of mb03 and mb04. the structural framework in this part of the section may be characterised as an antiformal stack, including thrust sheets mb03 below and kr01 above mb04. sedimentary units the martørv bakker section is dominated by the lønstrup klint formation. however, the most important sedimentological feature in the section is related to fig. 63. hydrodynamic brecciation in the lønstrup klint formation in the northern part of the martørv bakker section. photograph: june 1993. fig. 64. slump-fold structures formed in the thin-bedded sand layers enveloped in dark muds of the diamict sediments in the piggyback basin in the martørv bakker section. photograph: october 1998. 81 the basin developed at the top of the hanging-wall block connected to the normal fault displacing the mb02 thrust sheet (fig. 62). the deposits in this basin are regarded as an exotic part of the rubjerg knude formation, and are described under this heading below. the southern part of the martørv bakker section is unconformably overlain by the vendsyssel formation, the initially glaciotectonic truncation being superimposed by a post-glacial erosional unconformity. a holocene erosional unconformity truncates the vendsyssel formation as well as the glaciotectonic unconformity, and the peat deposited on this unconformity is up to 2 m thick in the northern part of the section, where it is covered by modern aeolian dunes up to 20 m high. lønstrup klint formation the lower part of the lønstrup klint formation is mudrich. it is exposed in the northern part of the section, where it was thrust above the hanging-wall flat from the décollement surface 25–28 m below sea level. the upper part of the formation is dominated by 0.5–1.5 m thick sand beds interlayered with thin beds of horizontally laminated mud, which typically has been mobilised to form hydrodynamic breccias. in the southern part of the section, the beds are strongly affected by ball-and-pillow deformation (fig. 63). rubjerg knude formation the rubjerg knude formation comprises two units: (1) a lower 3 m thick sand unit only exposed along the prominent normal fault in the central part of the martørv bakker section, and (2) a c. 15 m thick diamictite interpreted as a glaciolacustrine mud with redeposited clasts. the first unit was deposited on the l/r-unconformity, and comprises light yellowish medium-grained sand (fig. 26). the diamictite unit rests partly on the lower sand unit, and partly on the l/runconformity at the top of the lønstrup klint formation. the diamictite is dark grey, and comprises weakly laminated mud interbedded with structureless, irregularly distributed matrix-supported beds containing fig. 65. a schematic diagram explaining the development of the hydrodynamic brecciation displayed in fig. 63. the formation of the structure was the result of three phases of deformation. in the first phase, a succession of sandy turbidites interbedded with mud (1) was affected by loading to form the ball-and-pillow structures (2). in the second phase, the ball-and-pillow structures were displaced by thrust faulting (3). during the third phase, the mobilised mud intruded up thorough the thrust-fault surface (4), demonstrating the syntectonic development of the hydrodynamic brecciation. 82 unsorted clasts in a matrix of sandy mud (fig. 26). in the lowermost 2 m of the diamictite succession, the matrix-supported clasts were probably derived by redeposition of coarse-grained material eroded from the l/r-unconformity to the north. about 3 m up in the succession, a c. 1 m thick bed occurs with clay clasts 10 cm in size deposited in a weakly clay-laminated and sand-streaked silty mud. this bed is overlain by three 1.5–3 m thick units of isoclinally slump-folded, thin-bedded, fine-grained sand encased in dark structureless mud (fig. 26). these units are interpreted as slump-folded sheets derived from the lønstrup klint formation (fig. 64). the middle and upper slump-units are separated by a c. 5 m thick interval dominated by sandy mud with scattered clasts and a few thin sand beds. the uppermost sand bed was not af fected by slumping and shows large-scale trough cross-bedding. the lower part of the diamictite succession is strongly disturbed by hydrodynamic brecciation with mesoscopic-scale diapirs rising from the top of the lønstrup klint formation and penetrating upwards into the diamict sediments. this indicates that the diamictites were part of the main sedimentation affected by glaciotectonic disturbances. structures in the martørv bakker section, three types of structural elements were studied: (1) imbricate duplexes dominating the southern part of the section, (2) the normal fault in the central part of the section, and 3) hydrodynamic brecciation contemporaneous with, or superimposed on, ball-and-pillow structures (figs 63, 65). imbricate duplexes the imbricate duplexes in the southern part of the section constitute rhomb-shaped segments 10 to 25 m in size bounded by low-angle thrust faults. minor imbricates may occur along the thrust faults, but the thrust faults are mainly narrow fracture surfaces without significant brecciation. a mini-scale example of duplex formation is shown in fig. 66. although the structure is related to intraformational deformation of the beds, it illustrates instructively the formation of hanging-wall ramp propagation of a stacked footwall ramp. thus the footwall ramp is formed by the trailing edges of two duplex segments that were displaced one over the other to form a single planar ramp for the propagation of the upper thrust sheet. flame-like upright minor anticlines are interpreted as compressed hanging-wall anticlines formed during the sequential propagation of the various ramps. on top of the upper footwall hinge, a radial flame structure probably indicates the site of incipient diapirism (fig. 66). the formation of foreland-dipping thrust structures above the tip of a lower duplex segment is also apparent. normal fault the normal fault in the central part of the section is an uneven fault plane striking e–w with a dip of about 45° to the south. displacement along the fault plane is about 25 m, and a set of minor normal listric faults displace the top of the hanging-wall block (fig. 62). the footwall block comprises the lønstrup klint formation, which decreases in thickness southwards and forms an irregularly folded drag along the fault plane. at the top of the hanging-wall block, the diamict sediments described above are discordantly superposed on the light-coloured sand at the base of the rubjerg knude formation. hydrodynamic brecciation ball-and-pillow structures occur in the sand-rich parts of the lønstrup klint formation, where hydrodynamic mud mobilisation created chaotic breccias (fig. 63). the initial size of the sand ball-and-pillow structures is related to the primary thickness of the beds, but subsequent to the sedimentary load deformation they were distorted and deformed during thrust-fault related mud remobilisation. in fig. 63, the distorted ball-andpillow structures can be seen to be displaced by minor thrust faults, and these thrust faults were intruded by mobilised mud. the sequential development of this hydrodynamic brecciation is illustrated in fig. 65, where three phases of deformation are recognised, although these probably developed progressively during thrust-fault displacement and related loading of superposed thrust sheets. interpretation of structural development the thickness of mb01 implies that the décollement surface in the southern part of the section is situated 83 at the 20 m level. the thickness of mb03 is 30 m, implying that the décollement level stepped down 10 m somewhere in the central part of the section. a lower footwall ramp and a corresponding hangingwall ramp must therefore be included in the balanced cross-section. the l/r-unconformity reference surface on top of the mb02 thrust sheet was about 20 m above present sea level prior to the normal fault displacement, which indicates that a duplex 20 m in thickness is situated below the trailing end of mb02. the lower footwall ramp responsible for the fault-propagation folding of the antiformal stack in the northern part of the section, estimated from the bend of the trailing ends of mb02 and mb03, must be situated below mb02. the structural model therefore suggests that a subsurface duplex was formed by segments of the mb02 thrust sheet situated between the 20 and 30 m levels (mb02u1–mb02u3). the footwall ramp thus constitutes two 10 m thick duplex segments stacked on top of each other. consequently, a hanging-wall anticline with a foreland-dipping limb formed above the hanging-wall ramp of mb02 and was translated along a footwall flat. the model further suggests that the hanging-wall fig. 66. a model of duplex formation is here illustrated by a mini-scale structure related to intraformational deformation of beds in the lønstrup klint formation, central part of martørv bakker section. the duplex comprises two segments, which were derived from the bed underlying the lower footwall flat (lower fwf). the footwall ramp for the segments is situated to the left outside the frame of the figure. the lower segment is a relatively short one, which was thrust over by the upper segment during the push from the ramping of the upper thrust sheet. during propagation up the footwall ramp, the trailing edges of the two segments were displaced to form one planar ramp for the upper thrust sheet, which was further translated over the duplex to a foreland-dipping bend created above the tip of the lower duplex segment. the flame-like upright anticlines are interpreted as compressed hanging-wall anticlines formed during the sequential propagation of the various ramps. note the radial flame structures at the upper footwall hinge indicating incipient diapirism. the small normal faults to the left of the trowel (15 cm in size) are thought to reflect similar forelanddipping features in the subsurface. fwh, footwall hinge; fwr, footwall ramp; ufwf, upper footwall flat; hwf, hanging-wall flat; r, ramp. photograph: june 1993. 84 structure is a composite feature partly constructed by the hanging-wall anticline related to the tip of the mb02u3 segment folded over the footwall ramp of mb02u2, and partly by the hanging-wall anticline related to the translation of the main hanging-wall ramp of mb02 along the 10 m level. the foreland-dipping limb of this structure corresponds well with a 45° south-dipping normal fault with a vertical displacement of about 20 m. it is therefore concluded that the northern slope of the diamict sedimentary basin was formed by the normal fault reflecting the foreland-dipping limb of a hanging-wall anticline. the southern more gently dipping slope of the basin was formed by the bend of mb02 due to its propagation up along the footwall ramp and flat of mb01. this footwall thrust fault is a composite imbricate duplex, which hampers the exact distinction of ramp-flat relationships. the slumpfolded units in the basin are interpreted as the result of gravity slides derived from the crest of the hanging-wall anticline or the tip of the mb04 thrust sheet propagating from the north. the sediments filling the basin represent redeposited material derived from the thrust-fault elevated part of the lønstrup klint formation, the coarse-grained clastics on the l/r-unconformity, and the lowermost part of the rubjerg knude formation. the basin is interpreted as a piggyback basin with syntectonic deposition during the translation of the mb02 thrust sheet. kramrende section from the south, the first significant macroscopic-scale diapir occurs in the kramrende section (the kramrende diapir). although mobilisation features also occur in sections farther to the south, this diapir is regarded as the most distal in the glaciotectonic thrustfault complex. the thickness of the thrust sheet hosting the kramrende diapir suggests it is related to the fig. 67. the northern part of the kr01 thrust sheet where the lithostratigraphic reference section of the lønstrup klint formation (fig. 21) is situated. the thrust sheet is fault-bend-folded up along an initially low-angle (c. 8°) footwall ramp, which was subsequently folded into the present more steeply dipping orientation. note the reverse faults interpreted as small back-thrust faults. photograph: june 1993. 85 deep level of thrust-fault rooting, which is about 30 m below the l/r-unconformity. the thrust sheet to the south of the kramrende diapir and two thrust sheets to the north are included in the section because they are all affected by the structures related to the diapir. the frontal edge of the kramrende section is formed by the footwall ramp beneath the first thrust sheet (kr01, see plates 1, 2). this thrust fault is identical with the trailing-edge footwall ramp of the martørv bakker section, which is responsible for the marked monoclinal fault-bend folding of the kr01 thrust sheet (fig. 67). the steps to the beach are situated in the gully between the kr01 thrust sheet and the kramrende diapir. the steps lead up to the summerhouse area at oddervej, and are referred to as the kramrende steps or the oddervej trappe. tectonic architecture the kramrende section consists of four thrust sheets (kr01–kr04; plates 1, 2). the frontal thrust sheet (kr01) is ramped over the mb04 footwall ramp in the martørv bakker trailing thrust sheet. at the north end of kr01, the l/r-unconformity is situated about c. 10 m a.s.l., which indicates that the kr01 hanging-wall ramp propagated along an intermediate footwall flat (fwf). the upper ramping along the mb04 footwall ramp is fig. 68. ball-and-pillow structure developed in the sandy turbidite bed between 4 and 5 m in fig. 21. the structure is interpreted as a load structure formed immediately after sedimentation. additional load structures can be seen at the base of the sand bed, where flame structures related to the underlying clayey bed intrude the base of the sand bed. above the balland-pillow structure, pinch and swell structures within the sand bed are also interpreted as gravity load structures. photograph: june 1993. fig. 69. in the upper part of the kr04 thrust sheet in the kramrende section, the rubjerg knude formation forms a piggyback basin, which is overthrust by the br01 thrust sheet in the southern part of the brede rende section. the thrust fault displayed in the photograph is a hanging-wall flat for the thrust sheet br01 (br01hwf) and footwall ramp of thrust sheet kr04 (kr04fwr). a minor satellite thrust fault was formed below the main thrust at a late stage of fault propagation after the sand of the rubjerg knude formation had been somewhat compacted. photograph: june 1984. 86 responsible for the fault-bend-fold appearance of the kr01 thrust sheet. above the frontal part of the mb04 footwall ramp and flat, the kr01 thrust sheet is folded into a flat-topped anticline. in the involute part of this anticline, a splint or horse is present (the kr01 splint). this is a small thrust-sheet wedge torn off during thrust propagation, which created peculiar anticlinal features in the structural profile. the kr02 thrust sheet takes the form of a major diapir. initially the diapir was a thrust sheet that was displaced up along the kr01 footwall ramp and above the back of the kr01 thrust sheet. the kr03 and kr04 thrust sheets situated on the back of kr02 are characterised by marked dif ferences in the thickness of the rubjerg knude formation. in the kr03 thrust sheet, the thickness is only about 8 m, whereas in kr04 the thickness of the rubjerg knude formation is up to 20 m. this indicates that the kr04 thrust sheet was thrust over the upper footwall flat of kr03 at an earlier stage compared to a probably longer time of deposition in the piggyback basin of kr04. the c. 20° northerly dipping inclination of the kr03 footwall flat and related parallel structures is due to the bend caused by the propagation of kr02 along the footwall ramp. the main décollement level below the kramrende section is situated at the 30 m level. sedimentary units the description of the sedimentary units in the kramrende section is mainly based on sedimentological fig. 70. the kr01 footwall syncline developed below the kr01 footwall ramp (kr01fwr), which is overlain by the kr02 hanging-wall flat (kr02hwf). note how the mobilised mud migrated from the steeply dipping limb of the footwall syncline up into the kramrende diapir, intrusively penetrating the thrust fault. photograph: june 1995. fig. 71. mobilised mud from the lower part of the lønstrup klint formation in the kr02 thrust sheet intruded the turbidite sand beds in the upper part of the formation. photograph: june 1995. 87 logging of the lønstrup klint formation in the kr01 thrust sheet (fig. 21). the detailed section of the rubjerg knude formation at the top of this log is uncertain due to poor exposure and dif ficulty of access. the rubjerg knude formation exhibits variations in thickness throughout the kramrende section, and the description herein is based on scattered observations. a c. 1 m thick homogeneous, structureless sandy till caps the kramrende section. no preferred clast fabric has been recognised in the till, and its stratigraphic position is uncertain, although the occurrence of rare rhomb porphyry erratics may indicate an affinity with the norwegian ice (kattegat till formation). lønstrup klint formation in the kramrende section, the lower part of the lønstrup klint formation mainly occurs in the kramrende diapir, within the kr02–03 thrust sheets. remobilisation of the mud has obliterated primary sedimentary structures, and the diapirism also affected the thrust sheet kr03, so that only primary bedding is recognisable in the uppermost part of the formation. the primary sedimentary structures of the upper part of the lønstrup klint formation are reasonably preserved in the kr01 thrust sheet (fig. 20), although the sediments here, dominantly constituting sandy turbidite beds (10– 50 cm thick), are strongly affected by hydrodynamic brecciation creating ball-and-pillow structures (fig. 68). the top of the lønstrup klint formation in the kramrende section is truncated by the l/r-unconformity, which displays an erosional relief of 1–2 m. rubjerg knude formation the rubjerg knude formation varies in thickness from only 6–8 m at the top of the kr03 thrust sheet to about 20 m in the upper part of the kr04 thrust sheet (it is absent in kr02). the main part of the formation comprises thick-bedded, large-scale cross-bedded, medium-grained light yellow-grey sand. some beds are rich in heavy mineral sand, which occurs in parallel-laminated strata. in the kr04 thrust sheet, the heavy mineral beds are present about 6 m above the l/runconformity and again about15 m above the base. these beds have a characteristic content of small (0.1– 1 cm), grey clayey mud-clasts, which are interpreted to reflect erosion of muddy thrust sheet units in the vicinity of the depocentre. pedersen (1987) described deposits, referred here to the rubjerg knude formation, that accumulated syntectonically in footwall growth synclines. the deposits were characterised as ‘banana’ shaped basins, and a similar type of sedimentary/structural feature occurs in the kr04 thrust sheet (fig. 69). the rubjerg knude formation at the top of the kramrende section can be regarded as a piggyback basin and the footwall syncline as a growth-fault syncline. at the top of the piggyback basin, large-scale trough cross-bedded sand is truncated by small satellite thrust faults (similar to that shown in fig. 69), which are truncated by superposed c. 1 m thick trough cross-bedded sand beds. three succeeding developments of this interference between thrusting and deposition reflect the syntectonic depositional dynamics of the piggyback basin. structures structures of significance in the kramrende section are described under the following headings: (1) thrust faults, and in particular associated footwall ramps and footwall synclines, (2) the kramrende diapir, with the diapiric breccias and intrusive structures formed by mobilised mud, and (3) reverse faults, here interpreted as back-thrust faults. thrust faults the kr01 thrust sheet is bounded by the hangingwall ramp and flat (kr01hwr and hwf) at the base, and the kr01 footwall ramp and flat at the top (kr01fwr and fwf). the wedge-shaped geometry of the kr01 tip implies that the kr01hwr had a low angle of inclination, dipping about 15°n. towards the trailing end of the thrust sheet, the ramp passes into a hanging-wall flat which is parallel to bedding in the lønstrup klint formation. this thrust fault now dips at 25°n, although it is a hanging-wall flat situated on a footwall flat. this is due to the fault-bend folding related to the thrusting in the trailing part of the martørv bakker section. at its trailing end, the kr01 thrust sheet is folded into a footwall syncline (fig. 70). the bend of the northern limb in the syncline lifted the l/r-unconformity up to a position nearly 5 m higher than in the horizontal involute part of the fold, and the bedding in the lønstrup klint formation was tilted into a nearly vertical position (fig. 70). 88 the kr02 thrust sheet was thrust up along a c. 25° dipping footwall ramp (kr01fwr) onto the upper footwall flat above the rubjerg knude formation of the kr01 thrust sheet. the kr02 thrust fault is apparently a hanging-wall flat which indicates a rather long displacement for thrusting. the top surface of kr02 is a footwall flat upon which the kr03 hanging-wall flat is situated, only bringing different stratigraphic levels of the lønstrup klint formation into contact. the dip of the thrust fault is parallel to the dip of the kr02– kr01 thrust fault. the footwall flat of the kr03 thrust sheet is overlain by a c. 80 m long hanging-wall flat of kr04. the lønstrup klint formation is only about 8– 10 m thick above the hanging-wall flat (kr04hwf), indicating a fairly long intermediate flat (at the 10 m level below the reference surface). at the trailing end of kr04, the thickness increases indicating the presence of a hanging-wall ramp at the base of the thrust sheet, and a footwall syncline similar to that described in kr01 is present. in the upper part of the kr04 thrust sheet, the sediments deposited in the footwall growth syncline became overturned along the northern limb during translation of the hanging-wall ramp of br01, as described above (fig. 69). kramrende diapir the kramrende diapir constitutes the main part of kr02. the diapirism also affected the trailing end of kr01 (fig. 70) as well as some parts of kr03. as shown in fig. 71, mud of the lower part of the lønstrup klint formation in kr02 became mobilised and intruded the overlying turbidite sand beds and also penetrated upwards into the overlying kr03 thrust sheet. figure 70 illustrates a thrust fault penetrated by intrusive mud at the footwall ramp of kr01. here, the mobilised mud from the steeply inclined northern limb of the footwall syncline intruded into the mud-breccia along and above the hanging-wall flat of kr02. a large part of the boundary between kr02 and kr03 was deformed in a similar way and the primary layering destroyed. reverse faults in the kramrende section, significant steeply dipping reverse faults occur in the northern part of kr01 (fig. 67), and in the middle part of the kr03 thrust sheet. the displacement is only about 30–50 cm on the steep south dipping faults in kr01; the spacing between the faults varies from 3–9 m, and often the faults can be traced down into the tectonic breccia above the fig. 72. reverse fault-splay fan developed on the back of kr03 and also displacing the overlying kr04 thrust sheet. the structure is interpreted as a back-thrust fault-splay formed during the kr03 propagation over the upper footwall ramp situated on the back of kr02. photograph: july 1994; figure at fault-splay centre for scale. 89 hanging-wall flat. in the kr03 thrust sheet, the reverse faults form a fault splay fan with individual faults dipping moderately to steeply to the south (fig. 72); displacement varies from about 20 cm up to about 3 m. the reverse faults appear to have formed during thrust-sheet propagation over a ramp hinge and are interpreted as back thrusts. a bend over a shallow dipping ramp will only result in minor displacement on steeply dipping back thrusts, whereas bending over steeply dipping ramps creates low-angle back thrusts with potentially greater displacements. interpretation of structural development the balanced cross-section indicates that the kr01 thrust sheet was about 350 m long, of which a major part of the front tip has been eroded away. the amount of displacement is deduced from a series of balanced approximations to be 160 m (see plate 2). the footwall ramp of kr01 is interpreted to root in the 30 m décollement level, which indicates that the lower hanging-wall ramp of kr01 was displaced onto the intermediate footwall flat above the trailing segment of mb04. thus the lift and steep tilt of the northern limb in the kr01 footwall syncline is interpreted to have formed during the displacement of the lower hanging-wall ramp (kr01hwr) along the trailing-end footwall flat of mb04 (mb04fwf). the tip of the kr02 thrust sheet has been eroded away. to avoid exaggeration, the thrust fault is interpreted to have continued only about 15 m up in the air further to the south, which implies that displacement of kr02 was in the order of 50 m. the main part of the kr02 thrust sheet present in the cross-section is above the hanging-wall flat brought up from the 30 m décollement level. it is evident that the mobilised mud was derived from this low level and that it was activated in diapirism during the thrust-sheet propagation over the footwall ramps (kr01fwr) and its hinge bend. it is difficult to estimate the displacement for the fig. 73. the brede rende diapir in the frontal part of the brede rende section. the frontal part of the br02 thrust sheet was thrust up on its hanging-wall ramp (br02hwr) along the footwall ramp, which turned into the kr04 footwall flat (kr04fwf). at an early stage of thrusting, probably while the hanging-wall ramp passed by a lower footwall hinge, diapirism developed. the mobilised mud also intruded the l/r-unconformity and formed mushroom-shaped diapirs in the rubjerg knude formation. photograph: june 1984. the regional distribution of zeolites in the basalts of the faroe islands and the significance of zeolites as palaeotemperature indicators 123 the regional distribution of zeolites in the basalts of the faroe islands and the significance of zeolites as palaeotemperature indicators ole jørgensen the first maps of the regional distribution of zeolites in the palaeogene basalt plateau of the faroe islands are presented. the zeolite zones (thomsonite-chabazite, analcite, mesolite, stilbite-heulandite, laumontite) continue below sea level and reach a depth of 2200 m in the lopra-1/1a well. below this level, a high temperature zone occurs characterised by prehnite and pumpellyite. the stilbite-heulandite zone is the dominant mineral zone on the northern island, vágar, the analcite and mesolite zones are the dominant ones on the southern islands of sandoy and suðuroy and the thomsonite-chabazite zone is dominant on the two northeastern islands of viðoy and borðoy. it is estimated that zeolitisation of the basalts took place at temperatures between about 40°c and 230°c. palaeogeothermal gradients are estimated to have been 66 ± 9°c/km in the lower basalt formation of the lopra area of suðuroy, the southernmost island, 63 ± 8°c/km in the middle basalt formation on the northernmost island of vágar and 56 ± 7°c/km in the upper basalt formation on the central island of sandoy. a linear extrapolation of the gradient from the lopra area places the palaeosurface of the basalt plateau near to the top of the lower basalt formation. on vágar, the palaeosurface was somewhere between 1700 m and 2020 m above the lower formation while the palaeosurface on sandoy was between 1550 m and 1924 m above the base of the upper formation. the overall distribution of zeolites reflects primarily variations in the maximum depth of burial of the basalt rather than differences in heat flow. the inferred thinning of the middle and upper basalt formation from the central to the southern part of the faroes is in general agreement with a northerly source area for these basalts, centred around the rift between the faroes and greenland. the regional zeolite distribution pattern is affected by local perturbations of the mineral zone boundaries that reflect local differences in the temperature, perhaps related to the circulation of water in the underground. the zonal distribution pattern suggests that these temperature anomalies are in part related to nw–se-trending eruption fissures or zones of weakness separating the present islands and are subparallel to transfer zones in the faroe–shetland basin. both the regional and the local distribution of zeolite assemblages are probably a reflection of the basic volcanic-tectonic pattern of the faroe islands. keywords: faroe islands, palaeogene basalt plateau, zeolite zone, palaeotemperature indicators _____________________________________________________________________________________________________________________________________________________________________________________ o.j., scandinavian asbestos & mineral analysis, kildeskovsvej 62, dk-2820 gentofte, denmark. e-mail: oj@oj-sama.dk © geus, 2006. geological survey of denmark and greenland bulletin 9, 123–156. available at: www.geus.dk/publications/bull geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19123 124 the zeolites of the faroe islands have been known for more than 300 years (debes 1673) although the islands remained nearly unknown to mineralogists until the end of the eighteenth century. because of increasing interest in mineralogy in the nineteenth century, the faroe islands were visited by many naturalists. one of these was brewster (1825), who proposed the name levyne for a new zeolite species he discovered at dalsnípa on sandoy. the first modern description of the distribution of faroese minerals was published by currie (1905), who visited the faroe islands and described the minerals at 120 localities. five years later a new description of the faroese zeolites was presented by görgey (1910). interest in the minerals of the faroe islands declined during the following 70 years until betz (1981) visited the islands and reviewed the classic localities. the present author started a literature study to discover, if a system of zeolite zones exists on the faroe islands similar to that described by walker (1960) in east iceland, but concluded that published descriptions were based on minerals from the same set of localities that were known to be rich in mineral species and where large crystals could be collected. this sampling bias meant that it was not possible to decide if zeolite zones existed on the faroe islands, so in 1979 a systematic mapping of the zeolites in the faroe islands was initiated by the present author. during the following 20 years, more than 800 localities were visited and about 3000 rock samples were investigated in the field and in the laboratory. the work was extended by studying samples from the vestmanna-1 and lopra-1 boreholes drilled in 1980 and 1981, respectively (jørgensen 1984; waagstein et al. 1984). in 1996, the lopra-1 borehole was deepened to a total of 3565 m (lopra-1/1a) and the secondary minerals in the deepened part of the lopra well were also described by the present author (jørgensen 1997). the aim of the present paper is to describe the secondary mineral distribution in the exposed parts of the faroe islands and in the lopra-1/1a and vestmanna-1 wells. the results of the mapping are used to estimate the palaeogeothermal gradients and the altitudes of the palaeosurfaces at various places in the faroes basalt succession. the following topics will be discussed: (1) the general conditions for the use of zeolites as palaeotemperature indicators and the statistical distribution of zeolites in a vertical profile, (2) the original thickness of the three basalt formations and the volcanic evolution of the faroese basalt complex, and (3) the regional distribution of the zeolite zones as a function of the thicknesses of the middle and the upper basalt formations. outline of the geology of the faroe islands the faroe islands (62°n, 7°w) have a total area of 1400 km2, an average height of 300 m above sea level and form part of the north atlantic brito-arctic cenozoic igneous province that extends from the british isles to greenland. the faroe islands consist almost exclusively of flood basalts that were erupted about 59–55 ma (waagstein 1988; larsen et al. 1999). the basalts on the exposed part of the faroe islands are divided into a lower, a middle and an upper basalt formation, separated by two horizons termed a and c in fig. 1 (see also fig. 4). according to rasmussen & noe-nygaard (1969, 1970) the volcanic evolution of the faroe islands may be summarised as follows: volcanic activity started west of the present islands with the eruption of the lower basalt formation. with time the production rate of lava slowed to a temporary standstill. during this quiet period, about 10 m of clay and coal bearing sediments were deposited (the a-horizon.). volcanic activity restarted with an explosive phase, resulting in the deposition of coarse volcanic ash and agglomerates. an effusive phase followed during which the middle basalt formation was erupted from sandoy nólsoy skúvoy stóra dímun lítla dímun streymoy kallsoy viðoy fugloy borðoy svinoy eysturoy 62°00'n 7°00'w 20 km0 koltur hestur vágar mykines upper basalt formation middle basalt formation lower basalt formation irregular instrusive bodies and sills coal-bearing sequence dykes suðuroy fig. 1. geological map of the faroe islands. from rasmussen & noe-nygaard (1969). geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19124 125 several vents and small fissures within the present group of islands. finally, volcanic activity moved farther east, away from the present islands, causing the lava flows of the upper basalt formation to transgress the middle basalt formation from the east. the discordant surface between the middle and upper basalt formations is named the chorizon. after the upper basalt formation was formed, the basalt plateau was intruded by dykes and sills. large sills were intruded near the boundary between middle and upper basalt formations in streymoy and eysturoy. tectonic activity continued long after the volcanism ended until the faroese basalt pile acquired its present gentle easterly dip. methodology sampling and mineral identification renewal of a large part of the road system of the faroes just before initiation of the fieldwork made it possible to collect samples in fresh road cuts and new quarries along the roads. after most of the road sites had been examined, the mountains were traversed and samples collected along the old paths between the villages. in addition to the samples collected by the author, the present investigation is based on 500 specimens of faroese zeolites collected privately by k. jørgensen and on the collection of faroese zeolites in the geological museum, copenhagen. the minerals were identified by their crystal morphology, optical properties, and x-ray diffraction (xrd) patterns or by chemical analysis carried out on a scanning electron microscope equipped with an energy dispersive analytical system. the xrd reference patterns were taken from gottardi & galli (1985). mapping of the mineral zones walker (1960, 1970) defined his zeolite zones by seven distinctive amygdale mineral assemblages. each zone was fig. 2. mineral temperature scale. the five zeolite zones are defined by the index minerals chabasite + thomsonite, analcite, mesolite, stilbite + heulandite and laumontite. the temperatures are shown at zone boundaries. abbreviations used for the various minerals are shown in table 1. an: analcite ap: apophyllite ca: calcite ce: celadonite ch: chabasite cl: chlorite cld: chalcedony ed: epidote ep: epistilbite ga: garronite gi: gismondine gy: gyrolite ha: harmotom he: heulandite la: laumontite table 1. abbreviations used for mineral names le: levyne me: mesolite: solid me*: mesolite: hair-like mo: mordenite na: natrolite ok: okemite op: opal ph: phillipsite pr: prehnite pu: pumpellyite qz: quartz sc: scolecite sm: smectites st*: stellerite † source: kristmannsdóttir & tómasson (1978), kristmannsdóttir (1982) and jakobsson & moore (1986) mineral zones approx. temperatures in °c † ch th* th an me* ph le me gy mo st he ep ap la pr pu ed cl ce sm qz cld op ca zeolite free zone 40–60 high temperature zone > 300 laumontite 190–230 stilbite-heullandite 110–130 mesolite 90–100 analcite chabazite-thomsonite 50–70 0 1 2 3 4 5 6 geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19125 126 defined by the presence or dominance of certain mineral species, termed index minerals, whose names are used to designate the zones. in addition to the index minerals, other minerals may be present as indicated in fig. 2. walker’s (1960) original zones are the carbonate, chabazitethomsonite, analcite, mesolite, laumontite, prehnite and epidote zones. the original mesolite zone was later subdivided into a mesolite and a stilbite-heulandite zone. this extended zone definition was adopted in the present investigation (fig. 2). the classification of the zeolite zones was normally based on the mineral assemblages of amygdales, and fracture fillings were used only in places without amygdales. classification of the mineral assemblages from samples from the lopra-1/1a and vestmanna-1 boreholes was originally based on the abundance of the individual index minerals expressed as the weight% of the total mass of index minerals (jørgensen 1984). in the present investigation, which is based on about 3000 samples, quantitative analysis was carried out only on two selected mineral assemblages, one from the middle and one from the upper basalt formation. the assemblages from most localities consist of a large number of minerals which makes it difficult to estimate the relative abundance of the different minerals. another complication was the fact that more than one index mineral often occurred at the same locality. the present investigation is therefore based on the first formed index minerals in the amygdales, i.e. the minerals that were deposited nearest to the host rock. where more than one first formed index mineral was present at a locality, the index mineral assumed to have the highest temperature of formation was chosen to map the zones. the mineral zones mapped in this way thus reflect the maximum temperature of mineralisation. this method is different from a mapping based on abundance of the minerals, which shows the distribution of the zeolite zones as the result of the main mineralisation. appendices a and b give the observed paragenesis in the 29 sections and two table 2. relative frequency (in %) of amygdales and mineralised fractures in the exposed part of the faroe islands mineral zeolites: analcite chabazite cowlesite epistilbite garronite gismondine heulandite laumontite levyne mesolite mordenite natrolite scolecite stellerite stilbite phillipsite thomsonite other minerals: apophyllite calcite celadonite chlorite csh gyrolite smectites silica minerals visited localities 32 67 1 <1 4 4 30 11 17 61 8 4 1 45 37 8 64 15 36 19 8 2 12 18 19 61 61 7 0 8 0 51 5 5 56 12 10 2 24 37 17 59 12 37 15 10 5 7 7 29 160 16 65 0 0 0 10 69 6 16 65 0 0 0 55 10 6 45 13 39 13 3 0 3 3 6 85 49 71 0 3 6 0 77 29 23 86 14 9 0 26 74 3 60 29 34 11 3 3 14 36 14 148 33 78 0 0 8 5 60 29 20 63 8 0 0 20 38 3 60 20 43 38 23 0 28 23 20 302 35 76 2 0 8 8 63 3 26 60 19 5 3 13 39 11 75 27 32 27 13 8 29 26 19 235 18 62 0 0 0 3 38 5 8 48 0 3 0 6 33 10 73 3 32 14 5 0 0 16 25 52 15 66 0 0 0 0 38 0 20 50 0 0 0 9 26 15 79 0 32 15 15 21 41 41 abundant minerals in bold face, common minerals in italics and rare minerals in normal type face. csh is calcium silicate hydrates. silica minerals are opal, chalcedony etc. average vágar suðuroy sandoy streymoy eysturoy borðoy viðoy geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19126 127 wells mapped. the most probable temperature range of deposition of the zeolite zones is indicated in fig. 2. the rules of zone classification stated above could not be followed strictly everywhere. in the southern part of suðuroy, mineralised vesicles are rare and, in this area, the mapping had to be based mainly on mineralised fractures. in the lopra-1/1a borehole, the study was based on cuttings. they included fragments of amygdales and mineralised fractures, and the first formed index mineral could be determined only when part of the host rock adhered to the sample. in order to establish a correlation between the zeolite zones and the temperature of formation of the minerals, the vertical distribution of index minerals and temperatures was examined in a number of boreholes in the geothermal areas of iceland (kristmannsdóttir & tómasson 1978; kristmannsdóttir 1982). the result was the mineral–temperature scale shown in fig. 2. examination showed that the temperatures at the boundaries of individual zeolite zones varied from place to place. this is probably caused by the fact that zeolites can be formed within a broad range of temperatures and that variations in the chemical composition of the rock and the hydrothermal solutions can affect the formation temperature of the zeolites (barth-wirsching & höller 1989; breck 1974). another problem that makes it difficult to determine accurately the palaeotemperatures at the zone boundaries is the fact that zone boundaries are not well defined lines, a problem that will be discussed below. the temperatures at the boundaries of the zeolite zones are therefore indicated in fig. 2 at the lowest and highest temperatures that occur at the icelandic zone boundaries. as mentioned above, the original classification of the mineral assemblages of the lopra-1/1a and vestmanna-1 drillholes was based on the most abundant index zeolite. it was therefore necessary to re-classify the mineral assemblages of the two drillholes according to the method used in the present investigation. this had a rather small effect on the zonation of the lopra-1/1a drillhole. however, the first formed minerals from the vestmanna-1 borehole are overgrown by abundant chabazite and thomsonite. neglecting these later deposits, changes the zonation from a simple chabazite-thomsonite zone to an alternation between the mesolite and stilbite-heulandite zones. the distribution of minerals, zones and temperatures frequency of occurrence of minerals table 2 shows the 17 zeolites and 8 associated minerals that were recorded in amygdales and mineralised fractures of the basalt in the islands of suðuroy, sandoy, vágar, streymoy, eysturoy, borðoy and viðoy. in addition to the minerals listed, prehnite, pumpellyite, native copper and pyrite were found in the lopra-1/1a borehole. all the zeolites listed in table 2 have been reported previously from the faroe islands with the exception of garronite, ch te le ph ap gy cl ca me st he la an su sa va st ey bo vi su sa va st ey bo vi su sa va st ey bo vi r el at iv e fr eq ue nc y as % o f l oc al it ie s r el at iv e fr eq ue nc y as % o f l oc al it ie s r el at iv e fr eq ue nc y as % o f l oc al it ie s 100 80 60 40 20 0 80 60 40 20 0 50 20 30 40 10 0 s n s n s n fig. 3. variation diagrams of the relative frequency of minerals on suðuroy (su), sandoy (sa), vágar (vá), streymoy (st), eysturoy (ey), borðoy (bo) and viðoy (vi). contractions for zeolite names are listed in table 1. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19127 128 c 3500 3000 2500 2000 1500 1000 500 0 up pe r ba sa lt fo rm at io n m id dl e ba sa lt fo rm at io n lo w er ba sa lt fo rm at io n su 1 su 2 su 3 su 9 va 1 va 4 va 7 w es tm an na -1 w el l st 6 w st 6 e sa 5 st 1 0 sa 1 sa 3 ey 1 ey 2 ey 3 ey 4 ey 8 ey 1 0 bo 1 bo 2 bo 3 v i 1 v i 2 st 2 w st 2 e th–ch zone an zone me zone st–he zone altitude of profile a c v i 3 fig. 4. stratigraphic location of the zeolite zones in the 28 sections through the exposed part of the faroe islands. the data shown here are tabulated in table 3. which was found for the first time at several localities during the mapping reported here. the minerals in table 2 were divided into three classes: (1) very frequent minerals that occur at 60% or more of the localities; (2) common minerals that occur at between 15% and 60% of the localities; (3) rare minerals that occur at less than 15% of the localities. the frequency of occurrence is given as percentages of localities examined on each island. from the column named average in table 2, it is seen that chabazite, mesolite and thomsonite are the most frequent secondary minerals within the exposed part of the faroe islands, followed by stilbite, stellerite, heulandite, analcite and calcite. epistilbite and scolecite are rare minerals in the exposed part of the faroe islands, but they are common in the lopra-1/1a well. regional distribution of minerals and zones fig. 3 shows how the relative frequency of a number of index minerals and associated minerals varies from island to island. for most of the minerals, the relative frequency decreases from south to north and from west to east, but for the minerals of the analcite and the chabazite-thomsonite zone, the relative frequency increases in the direction of viðoy (fig. 1). the variation in relative frequency of minerals reflects the regional shift in the distribution of mineral zones. the analcite and mesolite zones are the dominant mineral zones on suðuroy. from sandoy to vágar, the analcite zone is gradually replaced by a stilbiteheulandite zone that becomes widespread on vágar. on streymoy and eysturoy the stilbite-heulandite zone has a less widespread distribution, so that the stilbite-heulandite and the mesolite zones are of equal importance. the areal extent of the mesolite and stilbite-heulandite zones is further reduced on borðoy and viðoy, so the chabazitethomsonite zone becomes the major one on these two islands. description of mineral assemblages and zones figures 5–11 show the geographic distribution of the mineral zones on the islands suðuroy, sandoy, vágar, streymoy, borðoy and viðoy. in order to show the vertical distribution of the mineral zones, 28 local sections were constructed that are also shown on figs 5–11. the sections have been arranged such that it is possible to follow the changes in the mineral zones both geographically and stratigraphically (fig. 4). the data on which the sections are based are shown in appendices a and b, and table 3 gives the thickness of each zone. in contrast to the profiles drilled by the lopra-1/1a and vestmanna-1 boreholes, the profiles from the exposed part of the faroes have been constructed from observations along each section line, so the sections do not represent a single vertical profile through the lava pile. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19128 129 suðuroy (fig. 5) the distribution of the secondary minerals on suðuroy is remarkably heterogeneous, a feature noted by currie (1905). to the north of section su2 that extends across the island from fámjin to holmssund, nearly all vesicles and fractures of the basalts are mineralised, but to the south of the section, amygdales and mineralised fractures are rare. the boundary (section su2) between the two parts of suðuroy forms a transition zone in which the scattered vesicles are partly mineralised by an analcite assemblage composed of hair-like mesolite, thomsonite, analcite, chabazite, calcite, quartz and chalcedony. this mineral assemblage is found all along section su2, which means that no relationship exists between the secondary minerals and their stratigraphic position within the lava pile. to the north of section su2, the number of mineral species and the degree of mineralisation increases gradually northwards and the area just north of trongisvágsfjørður is in the mesolite zone. despite apparent regularity, the northern part of suðuroy is a mosaic of small areas in which the mineral assemblages vary from place to place. the largest area of this kind occurs around the summit of gluggarnir (443 m). at this locality, nearly all minerals listed in table 2 are present. the southern part of the island is partly devoid of zeolites. the most abundant minerals are quartz, calcite and chalcedony, while mesolite, thomsonite, chabazite, analcite, heulandite and stilbite are less frequent. the mode of mineralisation is also different on the two parts of the island. on the northern part of suðuroy, mineralised fractures and vesicles occur in equal numbers, whereas mineralised fractures are more common than amygdales on the southern part of the island, in spite of the fact that empty vesicles occur at many localities. because of this, section su1 is based mainly on mineralised fractures. despite the weak amount of mineralisation on southsu1+lo su2 su5 su9 vá1 vá4 vá7 st2w st2e vestmanna-1 st6w st6e st10 sa1 sa3 sa5 ey1 ey2 ey3 ey4 ey8 ey10 bo1 bo2 bo3 vi1 vi2 vi3 0 330 798 970 842 1241 1424 1461 1461 795 1799 1799 2344 2252 2545 2594 1831 1650 1745 2039 2080 2098 2084 2263 2091 2237 2140 2188 _ _ _ _ 350 250 219 _ 100 400 125 _ _ 50 _ _ 166 50 _ _ _ _ _ _ _ _ _ _ 530† – 480 373 300 180 463 325 272 200 385 500 340 – 200 280 270 350 300 230 300 180 503 310 150 100 _ 180 430 _ _ _ _ _ 82 50 _ 20 83 _ _ 200 116 _ 125 195 170 190 200 _ 182 190 190 208 200 _ _ _ _ _ _ _ 293 200 _ _ _ _ _ 46 _ _ _ 75 193 149 187 _ 70 415 251 375 400 table 3. stratigraphic position and thickness of mineral zones (see fig. 4) me zone† thickness in metres ch–th zone thickness in metres st-he zone thickness in metres * stratigraphic height within the exposed lava pile of the faroe islands. the mesolite zone is composed of the upper 330 m of the lopra-1 mesolite zone plus the 200 m thick mesolite zone of section su1. base level stratigraphic height in m* section an zone thickness in metres † geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19129 130 fig. 5. distribution of zeolite zones on suðuroy and on sections su1, su2, su5, su9 and in the lopra-1/1a borehole. 7°00'w n 62°00'n 7°00'w20 km 5 km thomsonite–chabasite zone analcite zone mesolite–scolecite zone stilbite–heulandite zone laumontite zone empty vesicles 1000 0 –1000 –2000 –3000 –4000 analcite zone mesolite zone stilbite–heulandite zone laumontite zone ht zone suðuroy sandvik hvalbiarfjørdur skálafjall su2 su5 tempilklettur frodbiarkambur gluggarnir fámjin trongisvágsfjørdur hólmssund hovsfjørdur vágsfjørdur lopra-1/1a spinarnir su1 lambaklettur m et re s ab ov e se a le ve l su9 su1 + lopra-1/1a su2 su9 su5 geus bulletin no 9 7 juli.pmd 07-07-2006, 15:15130 131 thomsonite–chabasite zone analcite zone mesolite–scolecite zone stilbite–heulandite zone laumontite zone empty vesicles skopun sandur skàlavik husàvik dalsnipa stórafjall hálsur skarvanes tyrilsválur søltuvik sandsvatn pætursfjall dalur sa1 sa3 sa5 1.3° 500 400 300 200 100 0 m et re s ab ov e se a le ve l th–ch zone an zone me zone st–he zone sandoy 7°00'w n 62°00'n 7°00'w 20 km sa5 sa1 sa3 5 km ern suðuroy, a clear zonation can nevertheless be discerned along section su1, where a 200 m mesolite zone is overlain by a 180 m analcite zone. the mesolite zone continues to a depth of 600 m below sea level in the lopra-1/1a borehole (fig. 5). the most abundant minerals there are mesolite, scolecite, stilbite, heulandite and mordenite. chlorite, mesolite and scolecite were deposited first. mesolite and scolecite are replaced by laumontite as the first formed mineral at a depth of –626 m, indicating the top of the stilbite-heulandite zone. at about –1200 m, epistilbite replaces stilbite as the first formed mineral so that the order of deposition becomes celadonite/chlorite-epistilbite-thomsonite-laumontite or celadonite/chlorite-epistilbite-laumontite-stilbite. since an epistilbite zone has not yet been defined elsewhere, it was decided to include the total interval between –1200 m and –2200 m depth in the laumontite zone. a high temperature assemblage of laumonite, mordenite-prehnite, pumpellyite, chlorite, calcite and quartz is found from about –2200 m to the bottom of the lopra-1a borehole at –3534 m. fig. 6. distribution of zeolite zones on sandoy and on sections sa1, sa3 and sa5. the arrow indicates the apparent dip of the zeolite zones on section sa3. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19131 132 800 600 400 200 0 me zone st–he zone thomsonite–chabasite zone analcite zone mesolite–scolecite zone stilbite–heulandite zone laumontite zone empty vesicles m et re s ab ov e se a le ve l eysturtindur akranesskarð vestmannasund hestfalsgjøgv oyragjøgv vá4 vá1 vá7 kvigan dalsá br eið á malinstindur sandavágur midvágur sørvágur høgafjall sørvágasfjørdur reyðastiggjatagi rógvukollur skjatlá 4° vágar 7°00'w n 62°00'n 7°00'w 20 km 5 km và1 và4 và7 the uppermost mineral zone preserved in the su1 section is part of an analcite zone. in east iceland, a chabazitethomsonite zone (walker 1960) and in east greenland a zeolite free zone (neuhoff et al. 1997) has been recorded in the uppermost parts of the basalt complexes whose total thicknesses are 700 m and 1400 m, respectively. if equivalent zones have ever existed on the southern part of suðuroy, they must have been considerably thinner than those on iceland and greenland, because the palaeosurface of the lower basalt formation was about 300 m above the ahorizon (see below). sandoy (fig. 6) the southern part of sandoy is strongly mineralised while the mineralisation in the northern part is weak. the area north of a line from søltuvík to skálavík is weakly mineralised by quartz, calcite, chalcedony, chabazite, hair-like mesolite, thomsonite and late formed stilbite or stellerite. nearly all fractures and vesicles are totally mineralised to the south of the line. the most abundant zeolites are analcite, chabazite, heulandite, mesolite, stilbite and thomsonite. no distinct boundary has been observed between the northern and the southern part of the island. sections sa1, sa3 and sa5 (fig. 6) show that four fig. 7. distribution of zeolite zones on vágar and on sections vá1, vá4 and vá7. the dip and strike of the zeolite zones is indicated on the map. geus bulletin no 9 7 juli.pmd 07-07-2006, 15:16132 133 zeolite zones exist on sandoy. in the area between sandur and søltuvik, the uppermost 50 m of a stilbite-heulandite zone are exposed. the stilbite-heulandite assemblage consists of chlorite, heulandite, stilbite, mordenite and occasionally of laumontite and apophyllite. the latter two minerals occur mostly in fractures in the basalt. however, the degree of mineralisation is low in the stilbite-heulandite zone and, by volume, only half of the vesicles are mineralised. the southern and the eastern parts of sandoy are dominated by a mesolite zone. the most abundant minerals are chabazite, heulandite, mesolite (solid or hair-like), stellerite, thomsonite and calcite together with minor gyrolite, gismondine and levyne. heulandite and mesolite are the first deposited minerals. dalsnípa at the south-east coast of sandoy is the type locality of levyne (brewster 1825). a detailed quantitative analysis of the mineral assemblage and the zeolite zones in section sa3 is given below. vágar (fig. 7) in contrast to suðuroy and sandoy, vágar is totally mineralised and the island may be divided into two areas. northeast of a line from sørvágur to sandavágur, a stilbite-heulandite assemblage occurs at 40% of the localities that has not been observed farther south. south-east of the boundary line, the localities are dominated by a mesolite and analcite assemblage. around miðvágur hair-like mesolite with up to 100 mm long crystal needles can be found in larger fractures and cavities in the basalt. the distribution pattern of the mineral assemblages on vágar is controlled by the rise of the stilbite-heulandite zone towards the north-east. a calculation shows that the stilbite-heulandite zone dips from between 1–3°ssw to 0.6°ese, while the basalt flows dip 3–4°ese, i.e. the mineral zones are discordant to the lava stratification. a quantitative analysis of the mineral assemblage and the zeolite zones on section vá1 is given below. streymoy (fig. 8) streymoy can be divided into a northern, a central and a southern area. a stilbite-heulandite assemblage occurs along the coast from tjørnuvík to langasandur in the north. the observed minerals are stilbite, heulandite, thomsonite, compact mesolite, laumontite, gyrolite, okenite, tobermorite, apophyllite, celadonite and smectite. the mineral assemblage in the amygdales changes gradually towards the west. the hydrated calcium silicates, stilbite and laumontite, disappear from the amygdales although they still occur in fractures in the basalt. at saksun near the north-west coast, the stilbite-heulandite assemblage is replaced by a mesolite assemblage, characterised by solid mesolite, thomsonite, heulandite, chabazite, calcite and montmorillionite. gyrolite and stilbite are present, but only in fractures. the distribution pattern is reversed in the central part of streymoy (between langasandur, vestmanna, dalsnipa and kollafjørður). there the stilbiteheulandite assemblage occurs along the west coast from vestmanna to dalsnipa, while a mesolite assemblage is found along the east coast between langasandur and kollafjørður. only the mesolite assemblage is found in the southern part of streymoy. because of the differences in mineral distribution, sections st2 and st6 have been divided into two columns, showing the eastern and the western parts of the sections, respectively (fig. 8). the dip of the zeolite zones in the northern part of streymoy is 2°sw and 2°nne the central area. the distribution of secondary minerals is rather complex at many localities in northern and central streymoy and shows repetitive zoning, i.e. a regular repetition of two mineral zones. for example, on the path from saksun to the summit of borgin (643 m), a mesolite zone is first encountered, then a stilbite-heulandite zone, then, near the summit, a second mesolite zone. repetitive zoning has also been observed at loysingafjall (638 m), where the mesolite zone is overlain by a stilbite-heulandite zone, and along the main road between the villages of vestmanna and kvívík. between vestmanna and kvívík, zones occur within which the vesicles and fractures of the basalt flows are mineralised, either by heulandite, stilbite, mordenite plus minor laumontite, or by compact mesolite, heulandite, stilbite, thomsonite and chabazite. the widths of the zones are from a few hundred metres to about 1 km. repetitive zoning also exists in the vestmanna-1 borehole. the first classification of the vestmanna-1 mineral assemblages was based on the mass concentration of major index minerals. since chabazite and thomsonite are the most abundant minerals, the entire mineral assemblage was classified as a thomsonite-chabazite assemblage. during the reclassification based on the first formed minerals for the present work, it became apparent that the drilled lava succession contains a repetitive zoning of mesolite and stilbite-heulandite assemblages (fig. 8). it is unlikely that repetitive zoning is caused by vertical fluctuations of the geothermal gradient within such relatively short distances, but the repetitive zoning may reflect flows of water of different temperature that changed geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19133 134 streymoy 7°00'w n 62°00'n 7°00'w 20 km thomsonite–chabasite zone analcite zone mesolite–scolecite zone stilbite–heulandite zone laumontite zone empty vesicles tjørnuvik haldarsvik sundini saksun hvalvik hósvik kollafjørdur kaldbaksfjørdur st2e st2w langasandur tórshavn kvivik givrufelli langafjall borgin saksunardalur loysingfjall vestmanna-1 st6w bøllufjall hundsarabotnur sund hvitanes sundshalsurdalsnipa øksnagjogv kirkjubøur st10 2°3.4° th–ch zone an zone me zone st–he zone 700 600 500 400 300 200 100 0 m et re s ab ov e se a le ve l st–he zone 2 me zone 2 st–he zone 1 me zone 1 0 –100 –200 –300 –400 –500 –600 m et re s ab ov e se a le ve l vestmanna-1 st 2w st 2e st 6w st 6e st 10 4° 3.4° st 2e st 6w st 6e st 10 st 2w saksunardalur loysingfjall bøllufjall hundsarabolnur sundshálsur hvitanes øksnaglógv kirkjubøur vestmanna-1 vestmannasund sund 1.5° 2° 5 km givrufelli langafjall borgin 2° 2° fig. 8. distribution of zeolite zones on streymoy and on sections st2e, st2w, st6e, st6w, st10 and in the vestmanna-1 borehole. the dip and strike of the zeolite zones is indicated on the map. the arrows indicate the calculated strike of the zeolite zones between the sections. geus bulletin no 9 7 juli.pmd 07-07-2006, 15:16134 135 gøtuvik lambavik sun din i selatrað breiða morskarnes raktangi nes æðuvik rituvik runavik lambareiði stórafjall syðrugøta ritufjall heltnará sandfelli urðará glyvur inran svinár nordskáli litlafelli skerðingur elduvik funningur slættaratindureiði gjógv ey10 ey4 ey8 ey3 ey2 ey1 skálafjørdur th–ch zone an zone me zone st–he zone 700 600 500 400 300 200 100 0 m et re s ab ov e se a le ve l oyndarfjørdur fuglaflørdur funningsfjørdur oyri kolbanargjógv kambur 5 km 0.8° thomsonite–chabasite zone analcite zone mesolite–scolecite zone stilbite–heulandite zone laumontite zone empty vesicles eysturoy 7°00'w n 62°00'n 7°00'w 20 km ey1 ey2 ey3 ey4 ey10 ey8 1.7° 2° 0.4° 0.8° 2.8° 1.1° 2° 2° 2° 2° 4 ° 4° fig. 9. distribution of zeolite zones on eysturoy and on sections ey1, ey2, ey3, ey8 and ey10. the dip and strike of the zeolite zones is indicated on the map. the arrows indicate the calculated dip of the zeolite zones between the sections. locally the vertical distribution of temperature. alternative explanations for repetitive zoning are: (1) local variations in the chemical composition of the basalt, or: (2) mineralisation in an open and closed system, caused by variations in the percolation speed of the geothermal water (barth-wirsching & höller 1989; gottardi 1989). eysturoy (fig. 9) eysturoy may be divided into two for descriptive purposes. north of a line from norðskáli to fuglafjørður, the stilbite-heulandite assemblage occurs at 64% of localities. south of that line mineral assemblage occurs at only 12% of the localities (fig. 9). the actual strike and dip of the geus bulletin no 9 7 juli.pmd 07-07-2006, 15:17135 136 zeolite zones can be determined by combining the dip of the zeolite zones along the sections with the dips between them. on fig. 9 it can be seen that the strikes and dips change to follow the changes in mineralogy. north of a line from svinár/norðskáli to fuglafjørður, the zeolite zones dip about 2° towards the east. to the south of the line, the dips of the zones shift gradually from 4° to the sw to 2° to the s. this means that the mineral zones are discordant to the lava bedding on the northern part of eysturoy, but nearly concordant to it on the southern part of the island. borðoy (fig. 10) the mesolite, analcite and chabazite-thomsonite zones are the only ones exposed on borðoy. at klakkur (section bo1) the vesicles and fractures in the basalt are mineralised by heulandite, stilbite, mesolite (massive and hairlike), thomsonite (massive and hair-like), chabazite, levyne, phillipsite, montmorillionite and celadonite, but no clear relationship exists between the distribution of mineral and height in the lava pile. since analcite and mesolite are the first deposited minerals in most vesicles, the mineral assemblage of section bo1 was classified as a mesolite zone assemblage. the poor zoning west of borðoyavik suggests 800 600 400 200 0 th–ch zone an zone me zone m et re s ab ov e se a le ve l h vann asund h arald ssu n d bordoyavik høgahadd hálgafelli klakkur ánir strond húsadalur norðtoftir depil bo rð oy + v ið oy 2° 1.1° arnafjordur bo3 bo2 bo1 thomsonite–chabasite zone analcite zone mesolite–scolecite zone stilbite–heulandite zone laumontite zone empty vesicles borðoy 7°00'w n 62°00'n 7°00'w 20 km bo 3 bo 1 bo 2 5 km fig. 10. distribution of zeolite zones on borðoy and on sections bo1, bo2 and bo3. the arrow shows the strike of the zeolite zones between bo2 and bo3. the common dip and strike of the zeolite zones on borðoy and viðoy is shown in the upper right corner. geus bulletin no 9 7 juli.pmd 07-07-2006, 15:18136 137 altitude metres table 4. quantitative analysis of section sa3, sandoy. the table shows the number of observed and calculated amygdales containing the index mineral per 25 amygdales 0 25 50 75 100 120 150 180 200 210 220 243 270 320 340 360 375 380 400 420 446 18 18 10 16 0 4 6 2 0 2 1 0 1 1 1 0 0 0 0 0 0 ncalnobs me 0 0 0 0 0 0 2 3 7 14 15 22 22 10 5 0 1 1 0 0 0 0.0 0.0 0.02 0.1 0.3 0.8 1.4 5.9 5.6 2.0 0.7 0.5 0.1 0.0 0.0 ncalnobs me* 14 12 8 5 4 3 3 1 1 0 0 0 0 0 0 0 0 0 0 0 0 9.2 7.2 5.5 4.3 2.8 1.7 1.1 0.9 0.7 0.5 0.2 0.1 0.0 0.0 0.0 0.0 0.0 0.0 0.0 0.875 ncalnobs th 0 0 0 0 0 0 0 2 3 2 0 3 5 9 10 0 14 14 23 23 23 0.1 0.1 0.1 0.2 0.3 0.5 0.7 1.0 1.5 1.8 2.2 2.8 4.4 8.5 11.2 14.0 16.3 17.1 20.1 22.8 24.6 ncalnobs th* 14 0 11 8 0 8 5 5 4 5 0 3 3 2 1 0 0 0 1 0 0 12.8 11.5 10.2 9.2 8.2 7.4 6.3 5.2 4.5 4.2 3.8 3.3 2.5 1.5 1.1 0.9 0.7 0.7 0.4 0.3 0.3 ncalnobs an 0 0 1 1 0 2 3 0 0 0 6 6 6 8 13 0 0 0 16 0 20 1.1 1.3 1.5 1.8 2.1 2.4 2.7 3.6 4.2 4.4 4.8 5.4 6.7 9.3 11.3 12.0 13.2 13.8 15.4 17.2 19.1 ncalnobs ch † rn,u: correlation coefficient of the regression line inh versus u(nobs). see equation (2). ‡ sh: standard deviation on hcal. in m. see equation (3). 9.1 6.6 4.1 2.4 1.3 1.1 0.6 0.4 0.1 0.0 0.0 0.0 0.0 0.0 0.0 0.0 0.0 rn,u† sh ‡ _ _ 0.988 _ _ 0.996 _ _ _ _ 0.991 _ _ 0.982 _ _ 0.952 10.3 13.0 15.9 21.4 24.9 11.9 21.3 18.6 15.5 12.2 44 17 14 32 56 59 13.4 11.3 ncalnobsnobsnobsnobsnobs ncalncalncalncal ch th me he st table 5. quantiative analysis of section vá1, vágar. the table shows the number of observed and calculated amygdales containing the index mineral per 25 amygdales metres altitude 0 0 0 0 0 0 2 2 3 3 6 3 5 6 _ _ 0.2 0.3 0.6 0.7 1.2 1.2 1.4 1.7 2.2 2.4 2.7 3.5 4.8 5.3 0 0 0 3 13 9 15 0 19 18 19 21 0 23 _ _ 0.8 1.7 3.9 3.9 10.1 10.7 12.6 14.8 18.4 19.4 21.1 23.8 25.0 24.8 0 0 0 0 0 1 3 6 7 9 10 0 24 0 _ _ 0.0 0.1 0.3 0.4 1.9 2.2 3.1 4.5 7.6 8.8 11.0 16.1 22.0 23.5 4 9 24 23 15 18 10 7 3 2 1 0 0 0 _ _ 3.7 11.3 22.7 24.8 17.7 16.3 11.9 7.9 3.1 2.2 1.2 0.2 0.0 0.0 3 15 21 0 7 5 0 0 0 0 0 0 0 0 _ _ 3.7 14.3 25.0 23.2 6.7 5.4 2.6 1.1 0.2 0.1 0.0 0.0 0.0 0.0 † rn: correlation coefficient of the regression line inh versus u(nobs). see equation (2). ‡ sh: standard deviation on hcal. in m. see equation (3). 10 100 200 235 340 350 380 410 460 475 500 550 610 630 rn,u† sh‡ 0.987 0.972 0.851 0.837 0.871 58 27 25 26 97 geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19137 138 that the area has been affected by several generations of mineralisation. the zoning becomes distinct farther eastwards on sections bo2 and bo3. the exposed part of the mesolite zone is 500 m on bo1, but 300 m on bo2 and 150 m on bo3 (fig.10). the analcite and thomsonite-chabazite zones also appear at higher levels on sections bo2 and bo3. the mineral zones have apparent dips of about 1º towards the north and east. the analcite zone is about the same thickness on both bo2 and bo3, which suggests that the vertical displacement of the two zones reflects differences in altitude of the palaeosurface of the basalt plateau. the lava flows of borðoy dip 1.6º to the se, which means that the zeolite zones are discordant to the lava bedding. viðoy (fig.11) the distribution pattern of the secondary minerals on viðoy is a continuation of that on borðoy. the mineral zones are displaced downwards compared to those of borðoy with the result that only between 100 m and 200 m of the mesolite zone is exposed on sections vi1 and vi2. on section vi3 a lower analcite zone about 200 m thick is exposed, which is separated from the thomsonitechabazite zone by about 130 m of repetitive zoning. the chabazite-thomsonite zone is about 430 m thick, which is the maximum thickness recorded for that zone within the basalts exposed on the faroe islands. quantitative analysis of mineral distributions once the position and temperatures are known of the boundaries of the zeolite zones, the geothermal gradient h van n asu n d th–ch zone 3 an zone 3 th–ch zone 2 an zone 2 th–ch zone 1 an zone 1 me zone m et re s ab ov e se a le ve l 800 600 400 200 0 malinsfjall tunnafjall enni vl1 vl2 vl3 2.1° 2.1° 2° vi ðo y + bo rð oy thomsonite–chabasite zone analcite zone mesolite–scolecite zone stilbite–heulandite zone laumontite zone empty vesicles 2° vi3 vi2 vi1 vi d vi k 7°00'w n 62°00'n 7°00'w 20 km 5 km viðoy fig. 11. distribution of zeolite zones on viðoy and on sections vi1, vi2 and vi3. the dip and strike of the zeolite zones is indicated on the map. the arrows indicate the strike of the zeolite zones in the profiles. the common dip and strike of the zeolite zones on borðoy and viðoy is shown in the upper left corner. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19138 139 and the altitude of the palaeosurface of the basalt plateau can be estimated using least squares regression, assuming a linear palaeotemperature gradient. however, in order to make a reliable estimate, the regression must be based on three or more zone boundaries. this requirement is fulfilled only where the lopra-1/1a and su1 sections can be combined. elsewhere, the number of exposed zone boundaries is too small to calculate a geothermal gradient. in order to overcome this limitation, an attempt has been made to estimate the position of unexposed mineral boundaries from a detailed analysis of the distribution of the exposed mineral zones. quantitative analyses carried out on the mineral assemblages of the lopra-1/1a and vestmanna-1 boreholes showed that the relative vertical frequency of the minerals followed a skewed distribution with one or more maxima (jørgensen 1984). this observation has been used to extrapolate non exposed zone boundaries on section vá1 on vágar and sa3 on sandoy. the two sections were chosen because they are through the middle and the upper basalt formations respectively, the distribution of minerals in them is simple and nearly all rocks sampled contain a large number of well developed amygdales. in order to examine the relationship between the thickness of the zeolite zones and the distribution of index minerals, the number of amygdales containing a particular index mineral was recorded for 25 amygdales (see below). to ensure that the amygdales were selected randomly, all amygdales in the samples from each locality were numbered and the 25 amygdales for examination were selected by using a computerised random number generator. after the amygdales had been chosen, the first formed index mineral was determined and the total number of amygdales containing the same index mineral was recorded. these results are presented as columns nobs in tables 4 and 5. trials were carried out using different exponential distribution functions to find the best correlation between altitude and number of amygdales containing the same index minerals. the experiments showed that the best fit between the observed distribution and the calculated distribution was obtained by the log normal distribution function: ni,cal = n0 exp(–½[(lnh – lnh0) / a]2) (1) where: ni,cal is the calculated number of amygdales containing index mineral i. n0 is the total number of amygdales investigated; in this case n0 = 25. h is the altitude of the sample above sea level. h0 is the altitude where the distribution function attains its maximum value. a is a constant. by transforming equation (1) to a linear form and replacing ni,cal by ni,obs, we obtain: lnh = au + lnh0, where u = ± [2 (lnn0 – lnni,obs)]½ (2) the constants a and h0 can be determined by linear regression on u and lnh. the calculated number of amygdales (ni,cal) containing index mineral i is shown in tables 4 and 5. a t-test shows that u and lnh fit a straight line at the 1% confidence level. the distribution curves in figs 12 and 13 give only a best estimate for the height above sea level of the zone boundaries. to assess the degree of uncertainty of these estimates (see below) we calculate the standard deviation of the altitude (h) defined as: sh = [1/(n – 2) s (h – hc)2]½ (3) where: n is number of pairs (h, ni) along the distribution curve. h is the altitude of the sample above sea level. hc is the calculated altitude of a point on the distribution curves corresponding to ni amygdales that contain index mineral i. in equation (3), n is reduced by 2 because of the loss of two degrees of freedom by the least squares estimation of a and h0 in equation (1) (miller & freund 1977). when ni, h0 and a are known, hc can be calculated from equation (1). sh for the distribution curves is shown in tables 4 and 5 and a graphic representation of the calculated distributions of chabazite, thomsonite, mesolite, analcite, heulandite and stilbite on sections sa3 and vá1 is shown in figs 12 and 13. the shape of the curves suggests that a temperature range existed around the altitude h0 in which conditions were favourable for the formation of a particular zeolite. where h < h0, palaeotemperatures decreased away from h0 to where they became too low for the zeolite to form. where h > h0, palaeotemperatures increased away from h0. the zeolite that was most stable at h0 would have been formed in some interval below h0 but, at higher temperatures, formation of the first zeolite would gradually be inhibited and another one would have become stable. so a zeolite will most likely be found between the maximum slopes of its distribution curve versus height, which corresponds to a palaeotemperature interval. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19139 140 la zone st–he zone me zone th–ch zone zeolite free zonea n z on e la zone st–he zone me zone th–c h zone zeolite free zonea n z on e la zone st–he zone me zone th–ch zone zeolite free zonea n z on e ho analcite–chabazite ho ho 1000800600400200–20 0–400–600–800–1000–1200 0 3 6 9 12 15 18 21 24 solid mesolite–hair-like mesolite solid thomsonite–hair-like thomsonite ch (obs) ch (cal) an (obs) an (cal) n um be r of a m yg da le s co nt ai ni ng a n/ c h pe r 25 a m yg da le s altitude (m above sea level) 1000800600400200–200–400–600–800–1000–1200 0 3 6 9 12 15 18 21 24 altitude (m above sea level) n um be r of a m yg da le s co nt ai ni ng m e/ m e pe r 25 a m yg da le s 1000800600400200–200–400–600–800–1000–1200 0 3 6 9 12 15 18 21 24 altitude (m above sea level) n um be r of a m yg da le s co nt ai ni ng a n/ c h pe r 25 a m yg da le s ho ho me (obs) me (cal) me*(obs) me*(cal) th (obs) th (cal) th*(obs) th*(cal) la zone st–he zone me zone ch zone zeolite free zone la zone st–he zone me zone ch zone zeolite free zone chabasite–mesolite thomsonite–stilbite–heulandite 16001200 140010006004002000–200–400 800 3 6 9 12 15 18 21 24 altitude (m above sea level) n um be r of a m yg da le s co nt ai ni ng c h/ m e pe r 25 a m yg da le s 1000800600400200–200–400–600–800–1000–1200 0 3 6 9 12 15 18 21 24 altitude (m above sea level) n um be r of a m yg da le s co nt ai ni ng a n/ c h pe r 25 a m yg da le s ch (obs) ch (cal) me (obs) me (cal) th (obs) th (cal) he (obs) he (cal) st (obs) st (cal) + + + numbers are small. a zone boundary is defined for mapping purposes to be reproducible with a probability of 95%. the probability (p) of discovering ni,obs objects (amygdales containing the new zeolite i) among n0 objects (amygdales) can be calculated by means of the binomial distribution function (see e.g. kreyszig 1975 or miller & freund 1977). it can be shown that when n0 = 25 and p = 95%, then ni,obs = 3. fig. 12 shows that this reasoning can be applied to define the upper boundaries of the thomsonite-chabazite, analcite and mesolite zones at the heights where the upper end of the appropriate distribution curve intersects the line ni = 3. however, this method cannot be used on the upper boundaries of the stilbite-heulandite and laumontite zones, because stilbite, heulandite and laumontite do not exist as first formed index minerals in section sa3. fig. 12 shows that solid mesolite occurs most commonly in the mesolite zone and thomsonite occurs most commonly in the stilbite-heulandite plus the mesolite zone. if we assume that 95% of the two index minerals occurs within the zones in question, we can define the fig. 12. the calculated distribution of chabazite, analcite, thomsonite (compact and hair-like) and mesolite (compact and hairlike) on section sa3, sandoy. fig. 13. the calculated distribution of chabazite, mesolite, thomsonite, stilbite and heulandite on section vá1, váger. the calculated distribution curves shown in figs 12 and 13 show that the number of amygdales containing a particular zeolite decreases rapidly as |h – h0| increases. this has the practical consequence that it becomes harder to define a zone boundary by field mapping when sample geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19140 141 heights of the lower boundaries of the mesolite and the stilbite-heulandite zones where the lower end of the distribution curves interest the line ni = 3. that 95% of an index mineral occurs within the interval in question can be verified by plotting the accumulated distribution of the index mineral in a probability diagram. analcite occurs only sporadically below 340 m on section vá1 (fig. 13), so the analcite zone cannot be defined on this profile. the other zone boundaries were calculated as described above. estimation of palaeothermal gradients and altitudes of palaeosurfaces if we assume that the palaeothermal gradient was constant with depth, it may be estimated by linear regression on the data from the combined lopra-1/1a plus su1 sections, the mineral zone boundaries obtained by calculation and shown in figs 12 and 13 and the temperatures shown in fig. 2. the resulting estimates are shown in table 6. the new estimate from southern suðuroy is considered more accurate than that in jørgensen (1984), which was based on the mineral distribution in the lopra-1 borehole only. from these palaeogeothermal gradients and assuming a surface temperature of 7°c, the altitudes of the palaeosurface of the basalts has been estimated (table 6). the altitude differences between the estimated palaeosurface and stratigraphic marker horizons a and c is different at the three localities (table 7). on southern suðuroy, the palaeosurface was about 0.7 km (± 0.3 km) above present day sea level, i.e. close to the extrapolated top of the lower basalt formation. on vágar, the palaeosurface was 1.9 ± 0.2 km above the extrapolated top of the lower basalt formation (or 0.5 ± 0.2 km above the top of the middle formation), while on sandoy, the palaeosurface was 1.7 ± 0.2 km above the top of the middle basalt formation. this suggests that the focus of volcanism shifted laterally with time as will be discussed below. zeolite zone temperatures at the zone boundaries zeolite free 40– 60°c st–he 110– 130°c palaeothermal gradient °c/km altitude of palaeosurface m above sea level correlation coefficient r altitude of zone boundaries (m above sea level): lopra-1/1a + su1 vá1 sa3 0.9547 0.9181 0.8689 table 6. estimated palaeothermal gradients and the altitude of the palaeosurfaces at lopra-1/1a and sections su1, vá1 and sa3 ch–th 50– 70°c me 90– 100°c la 190– 220°c –590 380 425 –1200 – 825 –2200 – – 443–983 1760–2080 1150–1524 – 1325 653 – 900 150 66 ± 9 63 ± 8 56 ± 7– – table 7. estimated thicknesses of the basalt formations along various sections across the faroe islands altitude of palaeo surface in m above sea level (table 6) present stratigraphic thickness in m local altitude (m) of aand c-horizons altitude of palaeosurface in m above aor c-horizon 443–983 1760–2080 1150–1525 > 31001 14102 700–9001,3 a-horizon: a-horizon: c-horizon: 1550–1924 sources: 1) larsen et al. 1999, 2) waagstein & hald 1984, 3) waagstein 1988. s. suðuroy w. vágar e. sandoy area section formation lopra-1/1a + su1 vá1 sa3 l. formation m. formation u. formation a-horizon: 700 a-horizon: c-horizon: –400 –257–283 1700–202060 geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19141 142 discussion palaeothermometry and zone boundaries the use of zeolites as palaeotemperature indicators is based on the assumptions: (1) that the zeolite zones reflect univariant equilibrium with a high coefficient dp/dt of the clausius-clapeyron equation and: (2) that the formation temperature of the minerals is independent of the chemical composition of the hydrothermal solution. assumption (1) is fulfilled because the properties of condensed systems are nearly independent of pressure, if the pressure is not extremely high and the temperature is below the supercritical temperature (374ºc) of water. we can also estimate the error of the temperature determination when we ignore the external pressure of the system. the coefficient dp/dt is known only for a small number of systems that involve zeolites, but experimental studies of the systems stilbite-laumontite-h2o and laumontitewairakite-h2o show that dp/dt is of the order of 25–33 bars/degree in the range pfluid = 0–2000 bars (liou 1971; jové & hacker 1997). consequently, if the external pressure is equal to the fluid pressure and the original thickness of the lava pile is 2 km, the likely maximum error of the temperature determination is 8°c, disregarding the depth between the boundary of the zeolite zone and the unknown altitude of the palaeosurface. in contrast to the effect of pressure, the chemical composition of the rock and the hydrothermal fluids has a much larger effect on the formation temperature of the minerals. barth-wirsching & höller (1989) studied the formation of zeolites in glasses of different chemical compositions. they found that replacing rhyolitic glass by basaltic glass caused the formation temperature of different zeolites to increase by between 50°c and 100°c. this demonstrates the importance of choosing a reference area for the thermometry that consists of rocks with a chemical composition similar to that of the rocks in the area studied. other factors may also affect the formation of zeolites such as the texture of the rock, the content of glass and the porosity of the rock (gottardi 1989). the geothermal areas on iceland were used as a reference for the thermometry in the present study. the basalts from the faroe islands are all tholeiites, but show large compositional variations that range from picritic to ferrobasaltic (waagstein 1988). most of the icelandic basalts are also tholeiites, but minor amounts of acid rocks (rhyolites, andesites, granophyres, acid tuff ) are found, mainly associated with volcanic centres (sigurdsson 1967). this compositional variation of the icelandic rocks may partly explain the large variation in temperature at the boundaries of the zeolite zones mentioned above (fig. 2). from the description of the sections shown in figs 12 and 13 (and listed in appendix a), it can be seen that the distribution of index minerals varies gradually between successive zeolite zones. an index mineral that defines a zone may thus overlap the boundaries of neighbouring zones which makes it difficult to define the exact boundaries between mineral zones. the problem was solved by statistical analysis on sections sa3 and vá1 from which the boundaries of the zeolite zones could be defined as the locations where 3 out of 25 amygdales contain the appropriate index mineral. the overlap problem occurs in all the sections described in appendix a and, because of the lack of quantitative mineral data, the distributions in sa3 and vá1 were the only ones that could be described by a simple distribution model. on all other sections, the zone division was based on a crude estimate of the abundance of the index minerals around the zone boundaries. figs 12 and 13 show that the overlap between the zeolite zones varies from 100 m to 300 m. which means that, in the worst case, the zone boundaries could be determined with an accuracy of only ± 150 m when the zone boundary localities are based on a subjective estimate of the abundance of index minerals. volcanic and tectonic evolution the average of the three calculated palaeogeothermal gradients is about 60°c/km and there may be a small decrease in the gradient from the lower to the upper basalt formation. if real, this decrease could reflect either a reduction in heat flow with time or a variation in heat flow with locality. rasmussen & noe-nygaard’s (1969, 1970) summary of the volcanic evolution of the faroe islands that volcanic activity started in the west and moved eastwards with the times, must be modified, because evidence from the lopra-1 drillhole indicates that the lavas of the lower formation were erupted from local centres (waagstein 1988) and not from centres situated west of the present islands. this change might explain the change in the palaeogeothermal gradients shown in table 6. the eruption centres of the lower and middle formations were located in the faroe islands and the geothermal gradient was high. movement away to the east during eruption of the upper formation led to a decrease in the gradient. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19142 143 regional distribution of the zeolite zones at the time when the faroe lava pile was first mineralised, the thermal gradient seems to have been fairly constant, at least regionally and for some time. this is suggested by the rough equality of the calculated palaeogeothermal gradients from the southern, western and central part of the faroes that represent different stratigraphic levels (table 6). the overall regional distribution of zeolites is thus considered to reflect primarily variations in the maximum depth of burial of the basalt rather than differences in heat flow. the inferred palaeosurface on southern suðuroy is close to the extrapolated top of the lower formation (table 7), indicating that the total thickness of middle and upper formation lavas must have been small in this area. in contrast, in eastern sandoy about 50 km farther north where the exposed thickness of the upper formation is of the order of 1 km, the estimated palaeosurface is >1.5 km above the base of the upper formation or stratigraphically approximately 3 km above the lower–middle formation boundary. if we use the palaeogeothermal gradient calculated on sandoy, then the palaeosurface of the upper formation on viðoy may have been at about 1.3 km ± 0.2 km above sea level. these results suggest that the upper formation had a similar or only slightly smaller thickness in the north-eastern part of the faroes compared with the central part of the islands. on the other hand, the upper formation seems to have been much thinner or non existent in both the western and southern parts of the faroes (table 7) and the middle formation must also have been thin in the south. the inferred thinning of the middle and upper formations from the central to southern part of the faroes is consistent with a northerly source area for these basalts, centred on the rift between the faroes and greenland (waagstein 1988; hald & waagstein 1991; larsen et al. 1999). the thinning of the upper formation towards the west is consistent with rasmussen & noe-nygaard’s (1969, 1970) interpretation of an easterly source for this part of lava pile and may suggest a shift in the focus of volcanism. the first order regional zeolite distribution pattern is affected by local perturbations of the mineral zone boundaries (fig. 4). these perturbations show up as shifts in the dip of the zone boundaries within and between neighbouring islands as well as shifts in the degree of mineralisation. the latter effect is clearly seen towards the south. southern sandoy and northern suðuroy are heavily mineralised, although at different temperatures, whereas the vesicles of the basalt in the adjoining areas of northern sandoy and southern suðuroy usually contain no zeolites. on the northern and western islands, the zone distribution shows a tendency to symmetry around the narrow nw–se-trending sounds that separate the islands (figs 7–9). the distributions on the neighbouring north-eastern islands of borðoy and viðoy similarly seem to be mirror imaged, a distribution difficult to explain by variations in depth of burial. it is more likely that the distributions reflect local differences in palaeotemperature, perhaps related to the circulation of water underground with high temperatures in areas of up welling and low temperatures in areas of down welling. the symmetry of the zonal distribution patterns suggests that these temperature anomalies are in part related to nw–se-trending eruption fissures or zones of weakness separating the present islands (noe-nygaard 1968; rasmussen & noe-nygaard 1969, 1970). they are subparallel to the transfer zones in the faroe–shetland basin described by rumph et al. (1993) and later authors, and may indicate the presence of similar deep seated features. both the regional and the local distribution of zeolite assemblages probably reflect the basic volcanic-tectonic systems that led to the development of the faroe islands. acknowledgements i want to express my gratitude to the late arne noe-nygaard and jóannes rasmussen for discussions and support during the first phase of this project. i also want to express my thanks to the geological survey of denmark and greenland for financial support to the present project, to curator ole v. petersen, geological museum, copenhagen for permission to study the collection of zeolites from the faroes and to mrs. kitty jørgensen, næstved, who kindly made her collection of zeolites from the faroes available for my study. finally, i want to thank regin waagstein, james chalmers and kjeld alstrup for discussions and constructive criticism of the various versions of the manuscript. the comments of two anonymous reviewers are likewise greatly acknowledged. references barth-wirsching, u. & höller, h. 1989: experimental studies on zeolite formation conditions. the european journal of mineralogy 1, 498–506. betz, v. 1981: zeolites from iceland and the faroes. mineralogical record 12, 5–26. breck, d.w. 1974: zeolite molecular sieves: structure, chemistry and use. new york: john wiley. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19143 144 brewster d. 1825: a description of levyne, a new mineral species. edinburgh journal of science 2, 322–334. currie, j. 1905: the minerals of the faroes, arranged topographically. transactions of the edinburgh geological society (session 1905–1906) 9, 1–68. debes, l.j. 1673: færoæ et færoa reserata. hafniæ: suptibus daniels paaulli reg. bibl. (in latin and danish). görgey, r. 1910: ein beitrag zur topographischen mineralogie der färöer. neues jahrbuch der mineralogie und palaeontologie xxix, 269–315. gottardi, g. 1989: the genesis of zeolites. european journal of mineralogy 1, 479–487. gottardi, g. & galli, e. 1985: natural zeolites. in: wyllie, p.j., goresy, a.e., von engelhard, w. & hahn, t. (eds): minerals and rocks no. 18. berlin: springer verlag. hald, n. & waagstein, r. 1991: the dykes and sills of the early tertiary faroe islands basalt plateau. transactions of the royal society of edinburgh: earth sciences 82, 373–388. jørgensen, o. 1984: zeolite zones in the basaltic lavas of the faroe islands. annales societatis scientiarum faroensis. supplementum 9, 71–91. jørgensen, o. 1997: zeolites and other secondary minerals in cavities and veins, lopra-1/1a well, faroe islands, 1996, 8 pp. + plates. unpublished report. technical studies prepared for dansk olie og gasproduktion a/s 1997 (in archives of the geological survey of denmark and greenland, geus report file 26129). jove, c. & hacker, b.r. 1997: experimental investigation of laumontite ® wairakite + h2o; a model diagenetic reaction. american mineralogist 82, 781–789. kreyszig, e. 1975: advanced engineering mathematics, 707–709, 783–787. new york: john wiley. kristmannsdóttir, h. 1982: alteration in the irdp drillhole compared with other drillholes in iceland. journal of geophysical research 87 (b8), 6525–6531. kristmannsdóttir, h. & tómasson, j. 1978: zeolite zones in geothermal areas in iceland. in: sand, l.b. & mumpton, f.a. (eds): natural zeolites, occurrence, properties and use, 277–284. oxford: pergamon press. larsen, l.m., waagstein, r., pedersen, a.k. & storey, m. 1999: trans-atlantic correlation of palaeogene volcanic successions in the faroe islands and east greenland. journal of the geological society (london) 156, 1081–1095. liou, j.g. 1971: stilbite-laumontite equilibrium. contribution to mineralogy and petrology 31, 171–177. miller, i. & freund, j.e. 1977: probability and statistics for engineers, 50–63. englewood, new jersey: prentice-hall. neuhoff, p.s., watt, w.s., brid, d.k. & petersen, a.k. 1997: timing and structural relations of regional zeolite zones in basalts of the east greenland continental margin. geology 25, 803–806. noe-nygaard, a. 1968: on extrusion forms in plateau basalts; shield volcanoes of ‘scutulum’ type. science in iceland 1,10–13. rasmussen, j. & noe-nygaard, a. 1969: beskrivelse til geologisk kort over færøerne i målestok 1:50 000. danmarks geologiske undersøgelse i række 24, 370 pp. + map vol. (in danish with summaries in faroese and english). rasmussen, j. & noe-nygaard, a. 1970: geology of the faroe islands (pre-quaternary). danmarks geologiske undersøgelse i række 25, 142 pp. rumph, b., reaves, c.m., orange, v.g. & robinson, d.l. 1993: structuring and transfer zones in the faroe basin in a regional tectonic context. in: parker, j.r. (ed.): petroleum geology of northwest europe: proceedings of the 4th conference, 999– 1009. london: geological society. sigurdsson, h. 1967: the icelandic basalt plateau and the question of sial. in: björnsson, s. (ed.): iceland and mid-ocean ridges. societas scientiarium islandica xxxviii, 32–46. waagstein, r. 1988: structure, composition and age of the faroe basalt plateau. in: morton, a.c. & parson, l.m. (eds): early tertiary volcanism and the opening of the ne atlantic. geological society (london) special publication 29, 225–238. waagstein, r. & hald, n. 1984: structure and petrography of a 660 m lava sequence from the vestmanna-1 drillhole. in: berthelsen, o., noe-nygaard, a. & rasmussen, j. (eds): the deep drilling project 1980–1981 in the faroe islands. føroya fródskaparfelag, tórshavn, 39–65. walker, g.p.l. 1960: zeolite zones and dike distribution in relation to the structure of the basalts of eastern iceland. journal of geology 68, 515–528. walker, g.p.l. 1970: the distribution of amygdale minerals in mull and morvern (western scotland). in: murty, t.v.v.g.r.k. & rao, s. (eds): studies in earth sciences, west commemoration volume, 181–194. faridibad, india: today & tomorrow’s publishers. manuscipt received 3 july 2001; revision accepted 7 december 2001. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19144 145 appendix a: locations of sites along the sections discussed in the paper and minerals found in vesicles and fractures at each locality. the mineral zones are defi ned in fig. 2 profi le locality altitude (m) vesicles fractures zone su1 road exposure 0.5 km nor 25 no vesicles. he–me. he–st–ch. an–ch. ca. qz. 2–3 su1 the northern entrance of the sumba tunnel 70 – he–me. he–ch. me–an. ca. ch–an. 2–3 su1 road exposure at the w slope of siglidalur 200 – ch–an. ca. 1–3 su1 road exposure at small stream on the w slope of spinarnir 380 – ch–th*. op. qz. 1–2 su1 288 – ch–an. op. qz. ca. 1–2 su1 road exposure at stórá, 1km se of spinarnir 340 – th*–an. op. qz. ca. 1–2 su1 lambaklettur 235 – as above. 1–2 su2 road exposure 3 km e of øravík 100 empty vesicles. ca. qz. 1–2 su2 road exposure 2.5 km e of øravík 25 ce–an. ch–me*–th*. an. ca. st. 1–2 su2 exposure in dalsá, 1.5 km w of øravík 100 an, me*. ch. ca. cld. cld. qz. 1–2 su2 road exposure just north of høgiklovningur, 2.5 km s of øravík 250 ca–an. ph–th*. cld. me*. 1–2 su2 278 empty vesicles. ca. 1–2 su2 nw slope of nónfjall 360 th*–an–ch. th–th*–an–ch. 1–2 su2 the summit of nónfjall 427 an–me* ca. qz. an. 1–2 su2 road exposure 0.8 km s of the church in fámjin 80 ac. ch. ch. an. qz. 1–2 su2 road exposure 0.5 km s of the church in fámjin 20 an–th. th*–ch. an. st. me*. 1–2 su5 høvdatangi, fr 0–25 empty vesicles. empty fractures. – su5 skarvatangi 60 na–(th, an). me–th. ch. me–me*. me–st. me–an. st–ch. st*–st. 2–3 su5 exposure at the road fr 90 an–st. me–st. th–st. th–me. me*–ch. 2–3 su5 – do – 130 na. ch. me–sc. me–ch. th–ga–ch. na–st*–st. an–th. st–ch. 2–3 su5 – do – 140 an–th*–ch. an–me. st–me–st. me–la, ch. 2–3 su5 – do – 250 me–me*. th–me. an. he–st. op, cld, ca. 2–3 su5 summit of kambur 483 ce–he. me in large acicular crystals like scolecite. ce–me*, ce–an. he. 2–3 su9 hamranes and the southern entrance of the tunnel hvalba–sandvík 0–100 he–st. he–me–ch. an–me. me–me*– gy. me–gy+ap. he–st–ap. me*–la. th–ch; he–(st, la). he–ap. he–la. la–ca. la–gy. 2–3 su9 the southern slope of skálafjall 70 an–he. me–th. cld. as in the vesicles. 2–3 su9 – do – 120 as above. 2–3 su9 – do – 160 scolecite-like me. an–he. me–he. he–me–me*. ca. 2–3 su9 – do – 200 as above. 2–3 su9 – do – 240 me–st. he–st. he–ch. ca–(ch, le). 2–3 su9 – do – 260 me–me*–ch. th–ch. 2–3 geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19145 146 profi le locality altitude (m) vesicles fractures zone su9 the southern slope of skálafjall 290 th*–ch. me–me*–ch. ca. 1–3 su9 the summit of skálafjall 374 as above. 1–3 sa1 exposure at the cost line and at road cuttings in søltuvík 0–50 many empty vesicles. he. st. la. ca. la–st–ca. mo–he–st. ca. 2–4 sa1 road exposure at the road sandur–søltuvík, 1.5 km west of sandur 60 he–st. he+me–st. th–st. he+me–st. ca. 2–4 sa1 large quarry at lake sandvatn, 1 km north of sandur 10 no vesicles. st–la–ca. he. st. 2–4 sa1 sprutthol, sandsvágur bay 30 he–me. ch. ca. st. ca. 2–4 sa3 húsavíklið n of húsavík 0–75 he–me. ce–me. th–me. st. 2–3 sa3 road exposure 2 km w of skálavík 78 me–ch. gi. cld. he–st–ca. 2–3 sa3 road exposure at hálsur, 3 km w of skálavík 119 me–th. me–me*. he–th. st–st*. ap. 2–3 sa3 urðarklettar nw of húsavík 120 me–th–ch. an–me–th. 2–3 sa3 húsavíklið n of húsavík 150 me–th–me. an. he. ca. 2–3 sa3 exposure at gravaráin 140 me–th. he–st. he–ch. an. st. la. 2–3 sa3 urðatklettar, nw of húsavík 150 ce–me–th–me*, me–me*–gy. me*. 2–3 sa3 – do – 180 ca–me–me*. an–me. 2–3 sa3 húsavíklið, n of húsavík 180–200 me–th*. an–th–me*. 2–3 sa3 exposure at gravaráin 210–220 me–ca–.me*. an–me*. an–th*. ca–he–st. ca–ap. 1–2 sa3 exposure at stórá 243 me–me*. me–th*–ca. gi. ca–st. 1–2 sa3 summit of heiðafjall 266 he–me*. an–me*–an. th*. 1–2 sa3 exposure at stórá 320–340 me–an–me*. ch. th*. 1–2 sa3 skriðubakki 360–380 th*–ch. me*–ch. me*. 1–2 sa3 – do – 400–420 ch–th*. th*–ch. an. many empty vesicles. 0–1 sa3 the summit of pætursfjæll 447 ch–th*. many empty vesicles. 0–1 sa5 dalsnípa 150 me–an. me–ap–st. th–ap. ch–ap. th–ch. an–ch. le. he–st–ap. cld. 2–3 sa5 the s slope of skúvoyafjall, 0.6 km nw of dalsnipa 280 me–st, me–an, he. 2–3 sa5 the summit of skúvoyafjall 354 ce–ch. ce–th–th*. st. 1–3 sa5 road exposure at the end of the road dalur–skuvoyafjall 308 he–me. th–me*. ch–th*. 1–3 sa5 road exposure 2 km sw of dalur 260 th–ap. st–me–le. th–le. he–me–la. 2–3 sa5 dalur harbour 0–30 ch. ch–le. an–th–th*. he–me–st. ch. le. me–ap–st. me–me*. 2–3 sa5 road exposure at kinnartangi 100 ch. ch–th–ch. st–me–gy. gy–me. 2–3 sa5 the se slope of stórafjall 160 ch. an–me–me*. st–th. ch. st. he. gy. 2–3 sa5 – do – 220 me–me*–ch. th–me. 2–3 sa5 – do – 260 as above. 2–3 sa5 – do – 300 ph. gi–me*. th*–le. me–me*. 1–2 sa5 – do – 340 as above. ch–le. he–me–st–ap. 1–2 sa5 – do – 360 as above. he–le. 1–2 geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19146 147 profi le locality altitude (m) vesicles fractures zone sa5 the summit of stórafjall 396 ch+ph. gi–me*. me*–ch. th–ch. 1–2 sa5 the ne slope of stórafjall 310 ch. me*–ch. he. 1–2 sa5 trigonometric station on the ne slope of stórafjall 217 ch. an. me–ch. he–st. he–me–gy. 2–3 sa5 road exposure at tjarnaheyggjur 60 as above. me–ca. 2–3 vá1 reyðastiggajatangi 0–10 he–st, an. cl–st–ca–la. 3–4 vá1 gásadalur, exposure along the path gásadalur–rógvukollur 100 he–st, st–la. cl–st–la. 3–4 vá1 – do – 200 he, st. me, la, an. 3–4 vá1 gásadalur, exposure at the path gásadalur– rógvukollur 235 he–th. la, ca. 3–4 vá1 gásadalur. the pass between knúkarnir– neytaskarð 340 he, st, th, an, cl. 3–4 vá1 grunnadalur, exposure at the branching of small streams 380 cl–he–th, cl–he–st, cl–st–me. st–la–st. 3–4 vá1 rógvukollur , exposure at the w slope 380 he–me, me–ch, th–ch–th, mo–he, mo–ch ± cld. me, ca. 2–3 vá1 neytaskarð, exposure at the se slope 400–420 as above. 2–3 vá1 the summit of rógvukollur 464 he–me, me–he–me, th–me, th–ch, me–ch. 2–3 vá1 djúpidalur (the nw slope of eysturtindur), exposure at stream 470–480 he–me, he–ch, me–ch, th–ch. 2–3 vá1 500 he, me, th, ch. 2–3 vá1 grunnadalur, exposure at the end of small streams 550 th–me*–ch. la, ca. 2–3 vá1 djúpidalur (the nw slope of eysturtindur) 600–610 ce–th–th, ce–ch. st–me–la, ca. 2–3 vá1 the plateau between eysturtindur and akranesskarð 620–640 an, th, me, sm. 2–3 vá4 oyrargjógv ferry harbour and the path to sørvágur 0–136 st–st*. ep–st. la. th. ch. st. la. 3–4 vá4 large quarry 1 km w of sørvágur 10 cl–st–st*. st–la. mo–he. mo–gy. an–gy. me–me*–ch. me–ap±sm. he–st–st*. ep. la, me. 3–4 vá4 sjatlá, 1.5 km n of sørvágsvatn 45–60 cl–he–st. st–la. ap. 2–4 vá4 exposure at n end of small road from sørvágur, just w of sjatlá 114 cl–st–me. cl–an. cl–ch±sm. st–la. ap. 2–4 vá4 – do –, exposure at small tributary of skjatlá 150 cl–he–st. st–ap. cl–me–th. cl–me– ch. me–ap. me–ch. me–ap. an–ap. he–ch. 3–4 vá4 – do – 190 as above. 3–4 vá4 – do – 220 cl–st–st*. cl–he ± la. th. mo–st. 3–4 vá4 end of breiðá (oyrargjógv) 250 ce–me. he–me. 3–4 vá4 – do – 304 ce–me–th. ce–th–ch. gy–me. 2–3 vá4 kvígandalur, exposure at the se tributary of kvígandalsá 250 an. he–le. me–le. an–th–ch. 2–3 vá4 husadalur, exposure at the w? tributary of kirjuá 275 an. me. he. me–an. 2–3 geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19147 148 profi le locality altitude (m) vesicles fractures zone vá4 kvígandalur, exposure at the se tributary of kvígandalsá 300–365 an. he. le. ch. 2–3 vá4 the cross between the path oyrargjógv– sørvágur and sandavágur–slættanes 436 me. th. st. ch. 2–3 vá7 road exposure 3 km north of sandavágur 100–120 ce–he–st. ce–st–me. ce–he–th. ce–mo–st. ce–me–st. ce–mo–he. ce–he–st. ce–me*–ch. ce–ap. 2–4 vá7 the western slope of malinstindur 219 as above. 2–4 vá7 – do – 235 ce–he–me. ce–me–gy. st–ca. 2–3 vá7 – do – 280 ce–me–ch.±ca. ±sm. 2–3 vá7 – do – 345 as above. 2–3 vá7 – do – 386 empty vesicles. – vá7 – do – 410 ce–he–me. th–me. th–an. st. ch. 2–3 vá7 – do – 500 as above. 2–3 vá7 – do – 538 he. an. me. th. th*. cld. ca. he–me. me*. cld. 1–2 vá7 – do – 563 me–me*. th*–ph. th*–ch+le. 1–2 vá7 the summit of malinstindur 580 ce–me*. 1–2 vá7 – do – 620 he. me. ph. ch. th. ca. cld. 1–2 vá7 – do – 683 me–he. me*–ch. ph–ch. 1–2 vá7 – do – 690 he. ch. le. th. cld. 1–2 st2w the path saksun–haldarsvík: kvíggjarhamar, saksun 0–100 he. st. th. me. ch. st–ca–st. gy. tb. ok. 2–3 st2w the slope of the mountain between skipá and gellingará 150 he–ch–th. me–st. 2–3 st2w – do – 200 me. ca. 2–3 st2w – do – 250 th–me. he–me. me–ch. st. la. 2–3 st2w – do – 310 as above. 2–3 st2w – do – 325 me. me*. th*–ch. he–me. ok. gy. ca. 2–3 st2w – do – 355 me–me*. th–me*. th*–ch. 2–3 st2w – do – 360 ce–th*. he–me*–sm. 1–2 st2w – do – 380 an–th*. th*–le. me*–ca. 1–2 st2w – do – 407 ce–ph–ch. ce–th*–ch. 1–2 st2w – do – 430 th*–ch. 1–2 st2w – do – 460 empty vesicles. 0–1 st2w – do – 555 le–ch. th–th*. 0–1 st2w víkarskarð 600 ce. th. ch. cld. 0–1 st2w the ne slope of gívrufelli 650 as above. 0–1 st2w the summit of gívrufelli 701 as above. 0–1 st2e víkarnes n of haldarsvík 0–30 he–st. ca. ca–la. 3–4 st2e the se slope of fjallið 100 as above + th. gy. ok. to–gy. la. 3–4 st2e – do – 150 he. me. th. to–st. th–gy–ap. 2–3 st2e the summit of fjallið 180 he–me. he–th*. ca. 2–3 geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19148 149 profi le locality altitude (m) vesicles fractures zone st2e the path haldarsvík–saksun: exposure 0.5 km sw of haldarsvík 148 he–st. me. me–th–ch. me–an. ph–le–ch. 2–3 st2e – do –, exposure 1 km sw of haldarsvík 200 he–st. he–me. 2–3 st2e – do –, exposure at the svínstiáir tributaries of kluftá, 1 km sw of haldarsvík 230 me. th. he. st. ca. 2–3 st2e – do –, exposure 1.2 km sw of haldarsvík 250 he. me. mo. he. st. ap. ca. 2–3 st2e – do –, exposure 1.5 km sw of haldarsvík 280 me–th–me–gy–sm. 2–3 st2e – do –, exposure 1.8 km sw of haldarsvík 350 me–th. 2–3 st2e – do –, exposure 2 km sw of haldarsvík 360 me–th. me–ph–ch. 2–3 st2e – do –, exposure 2.3 km sw of haldarsvík 370 me–th. an–me–me*. 2–3 st2e – do –, exposure 2.5 km sw of haldarsvík 400 th*–ch. me–th*. ph–ch. ch + le. 1–2 st2e – do – 410 me–me*. ch–le–ch. 1–2 st2e the se slope of víkartindur 420 me–me*–le. ch–ph–sm. 1–2 st2e – do – 440 .me*. th–th*–ch. 1–2 st2e – do – 500 th*. ch.cld. 0–1 st2e – do – 540 as above. 0–1 st2e – do – 620 as above. 0–1 st6w road exposure at the main road kvívík– stykkið, 1.5 km e of kvívík 50 he. st. la. me. th. ap. he. st. ap. me*. th. ch. an. 2–4 st6w tunnel workplace at the village of leynar 50 me. st. la. 2–4 st6w exposure at leynarvatn along the old road tórshavn-vestmanna 60–125 he. st. wa. la. ce. cld. st. ca. he. st. st*. la. me*, th, le. 2–4 st6w the path leynarvatn–hósvík, exposure 0.3 km ne of leynarvatn 150–190 ce–he–st. ce–me–gy. cld. ph, gy, ap. 2–3 st6w – do –, 0.4 km ne of leynarvatn 210 ce–he–st. ce–me–th. ce–st–la. 2–3 st6w – do –, 0.5 km ne of leynarvatn 260 ce–me–th. ce–he–me. ce–he–th. me–ca. 2–3 st6w the path leynarvatn–hósvík, exposure 0.5 km ne of leynarvatn 300 as above. 2–3 st6w – do –, 0.6 km ne of leynarvatn 340 as above. 2–3 st6w á halsi, 1 km ne of leynarvatn 380 as above. 2–3 st6w – do –, 1.5 km ne of leynarvatn 463 ce–he–me–me*. ce–th–ch. 2–3 st6w – do –, 1.9 km ne of leynarvatn 500 as above. 2–3 st6w – do –, 2 km ne of leynarvatn 510 me–he–me*. me–me*. th*–ch. he–me. 2–3 st6w hósvíksskarð 520–530 th. th*. ph. ch. le. + ce. 2–3 st6w the sw slope of bøllufjall 550 as above. 1–2 st6w the summit of bøllufjall 584 as above. 1–2 st6w the sw slope of gívrufjall 530 as above. 1–2 st6e road exps. between við áir and hosvík 15 he–st*. he–st–ap. he–st–ch. he– th–ch. th–st. th–ch. ga–th. an–ch. an–ap. an–th. an–th–ch±ce. he–ch–st. th–gy+ap. ch–th*–ap. 2–3 st6e the path hosvík–leynar: smørdalsá 160–203 he–me. he–ch–th, me–gy. he–an–th. 2–3 geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19149 150 profi le locality altitude (m) vesicles fractures zone st6e the path hosvík–leynar: smørdalsá 240–260 ce–me–he. ce–th–me. cld. me–th–ap–gy. 2–3 st6e – do – 360 as above. ca–la. cld. 2–3 st6e the ne slope between bøllufjall and gívrufjall 436 me–th. he–me–ca. he–th. 2–3 st6e – do – 480 he–me–me*. th–ch. th–th*. 2–3 st6e – do – 500 th–th*–gy. me*–gy. 2–3 st6e the ne slope of bøllufjall 530 th–th*. th–ph, th*–ch. 1–2 st6e the summit of bøllufjall 584 as above. 1–2 st10 large quarry 0.5–1 km nw of sund, kalsbaksfjørður 15 me. me*. th. an. ph. gi. ap. 2–3 st10 exposure at sundá, 1.2 km s of sund 227 he. me. ca. 2–3 st10 – do –, 1.6 km sw of sund 265 me. th. st–he. 2–3 st10 quarry 0.5 km sw of lambafelli at the high road tórshavn–kollafjørður 340 ce–he. ce–ch–le. ce–an. ce–ch. ce–he–sm, ce–he–ch. ce–he–th–ch. 2–3 st10 road exps. 1 km w of sundshálsur along the high road tórshavn–kollafjørður 310 ce–he–me. ce–me–ch. ce–th–me. 2–3 st10 small quarry at end of road to the water reservoir of havnardalur 170 he. me. th. ch. 2–3 st10 road exposure at the road tórshavn– velbastaður, 0.5 km n of velbastaður 160 ce–me. ce–th–ch. 2–3 st10 exposure at the road tórshavn–velbastaður, just n of velbasta ur 123 me–me*. me–ch. me.th–ch. th–me– me*. cld, ca. 2–3 ey1 road exposure at the road eiði–norðskáli, 0.4–0.5 km se of eiði 60–100 ap–gy–me. me–ph. he–gy–me. he– me–ch. st–aå–sm. th–ap. th–st–ap. th–ch–sm. st–la–st*–ca. he–st– ch–sm. 3–4 ey1 localities on the road eiði–funningur: quarry in djúpidalur, 2 km east of eiði 150 st–gy. st–th–gy. th–ga–th. me–th. me–ch. st–la. 3–4 ey1 50 m long road cutting on w slope of slættaratindur, 3.5 km east of eiði 200–230 me–th*–gy. me–th*–ch. me–he–me. he–me. st–st*. la–me*–sm. he–ap. st. cld. qz. 2–3 ey1 road exposure 0.4–0.5 km e of eiðisskarð just at the n slope of vaðhorn 336–346 ce–me–th*. ce–me–ch. 2–3 ey1 the n slope of vaðhorn 410 as above. he–st–st*. mo–st. 2–3 ey1 – do – 435 as above. 2–3 ey1 small quarry at the road fork eiði, funningur, gjógv, 1 km west of funningur (165) 165 he. st. me. me*. ch. 2–4 ey1 exposure at the coast line at funningur 5–10 mo. he. st. me. +ce. st. an. me. cld. 3–4 ey2 exposure at the coast between stórá and marká 0–20 mo. he. st. st. qz. cld. ca. 3–4 ey2 the path svínár–funningur 30–40 he–me. he–st–gy. he–me–gy. 3–4 ey2 – do – 100 me–th–ch. th–gi. th–me*. an–ch. an–th–ch. th–gy–ap. me–ap–st. 2–3 ey2 – do – 212 as above. 2–3 ey2 – do – 280 me. th. an. ch. an–th*–gy. 2–3 ey2 – do – 346 an. me. me*. th. th*. th–ch–sm. 2–3 ey2 – do – 400 ce–th–th*–ch. ce–ch. th–th*–ch. 1–2 ey2 – do – 420 ch–an. ch.gi–th*. an–ph–ch. 1–2 geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19150 151 profi le locality altitude (m) vesicles fractures zone ey2 kvígandalsskarð 460 me*. th*. ch. gy. sm. 1–2 ey2 – do – 480 as above. 1–2 ey2 the e slope of skerðingur 500 as above. 1–2 ey2 – do – 525 as above. 1–2 ey2 – do – 450 ce–th*. ce–me–me*–ch. 1–2 ey2 – do – 430 as above. 2–3 ey2 – do – 415 ce–an. ce–qz. cld. 2–3 ey2 – do – 380 me. 2–3 ey2 kvígandalur 363 he–me. he–th. ey2 skipagjógv 180 me–th–th*. ch–gi–th. an–th–ap–gy–sm. cl. ca. 2–3 ey2 – do – 80 he. me. th. ca. 2–3 ey2 – do – 50 gy–he. gy–th–me. 3–4 ey2 skipagjósoyran 0–10 he. st. th. la. gy. he. me–la–st. 3–4 ey3 large quarry 1 km s of oyri 20–30 me–th–sm. me–th–ch–sm. me–gy– me*. cl–he–th–ap. 2–3 ey3 oyrargjógv 100 me–th, me–gy. 2–3 ey3 – do – 210 no vesicles. he. st. me. th. gy. 2–3 ey3 the path oyri–skálafjørður 251 me. th. ch. 2–3 ey3 – do – 300 he–me. th–me. th–ch. 2–3 ey3 – do – 340 an. me*–ch. 1–2 ey3 – do – 350 me*–ch. th–th*. 1–2 ey3 – do – 400 th–st–gy. ch. st. 1–2 ey3 – do – 426 an–th*. an–th*–ca. ch. 1–2 ey3 – do – 495 th*–ch+le. 1–2 ey3 the sw slope of sandfelli 527 ce–th*–ch. ce–cld. he. me. ca. cld. 0–1 ey3 – do – 545 as above. as above. 0–1 ey3 the summit of sandfelli 572 as above. as above. 0–1 ey3 the path on the s slope of skálafjall 440 me–th. he–me. 2–3 ey3 – do – 405 as above. 2–3 ey3 – do –, just at öksnagjógv 200 as above. 2–3 ey4 small quarry just n of morskarnes, about 1 km n of nesá 20 he–th–ch. me–th. he–st. he–ap. ca. 2–3 ey4 the w slope of neshagi, just e of the locality above 140 cl–me–th. he–th. 2–3 ey4 exposure at skotá 194 as above. 2–3 ey4 – do – 230 me–me*. ch. 1–2 ey4 – do – 320 th*. me–me*. ch. he–me. he–st. 1–2 ey4 the se slope of kambur: exposure between the source of skotá and urðará 380 me*. ch, le +ce. 1–2 geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19151 152 profi le locality altitude (m) vesicles fractures zone ey4 the se slope of kambur: exposure between the source of skotá and urðará 400 as above. 1–2 ey4 – do – 430 th*. ch. le. ce. 0–1 ey4 – do – 462 ch. many empty vesicles. 0–1 ey4 the path steffanstangi–kambur 480 as above. 0–1 ey4 – do – 500 th–ch–th*. cld. ca. 0–1 ey4 – do – 540 as above. 0–1 ey4 the summit of kambur (trigonometric station) 593 as above. 0–1 ey4 the e slope of heygshagi 440 me*–ch. th*–ch. 1–2 ey4 – do – 400 as above. 1–2 ey4 – do – 250 as above. 1–2 ey4 markrá 250 me–me*. he–th. he–me*. 2–3 ey4 – do – 160 as above. 2–3 ey8 road exposure at the old road lervík– fuglafjørður/norðragøta, about 2 km nw of lervík 80 ce–mo–he–me. ce–ch. he–me–st. 2–3 ey8 localities along kálvadalsá in kálvadalur 200 mo–he–me–me*. an–th–ch. le–ch–le. he–me–st–st*. an–th. 2–3 ey8 – do – 270 as above. 2–3 ey8 – do – 300 he–me*–ch. an–th*–ch. 2–3 ey8 mannsgjógv 400 an–ph–ch. an–th–ch, me*–an. 1–2 ey8 the ne slope of navirnar 300 he–me–st*. he–ch, he–le. an–he. he–st. he–me. 2–3 ey8 the e slope of ritafjall 440 he–me*–ch. an–th*–ch. 1–2 ey8 – do – 490 he–th–ch. me*–st*. 1–2 ey8 – do – 520 th–ch. me*–ch. he. st. ca. 0–1 ey8 the summit of ritafjall 560 th.ch. le. 0–1 ey8 – do – 641 th*. nearly all vesicles are empty. 0–1 ey10 large quarry just n of the road fork skálafjørður–runavík–lambi 60–80 me–st*. th–st*. th–me–an. th–ch–st*. co–th. ha. he–st*. he–th–ca. ph–ch–ph. ha–ca. 2–3 ey10 the sw slope of ritafelli, ne of the locality above 180 no vesicles. he–me–ch–sm. me*–le–ch–sm. an–me*–sm. 2–3 ey10 – do – 200 me–me*. ch. th*. ph. an–me*–sm. me*–ph– le. ap–sm. 1–2 ey10 – do – 230–240 as above. 1–2 ey10 – do – 270 he–th–ch. th*–ph. an–th*, ch–st. an–st. ap–th–ph. ca. st–la. me–ap. 1–2 ey10 the edge of ritafelli 350 st, me*. th*. ph. ch. ce. 1–2 ey10 the sw slope of stórafjall 380 th*. ch. ph. 1–2 ey10 the w edge of stórafjall 440 th. ph. ch. an. ce. an. st. ch. ph. sm. 0–1 ey10 – do – 490 empty vesicles. st, th*. ch. ca. cld. sm. 0–1 geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19152 153 profi le locality altitude (m) vesicles fractures zone ey10 the w edge of stórafjall 520 th. many empty vesicles. 0–1 ey10 the summit of stórafjall 567 as above. 0–1 bo1 large quarries at klakkur 60 me–ch. he–me. cld. mo–st–la. ca–st. 2–3 bo1 the ne slope of klakkur 100 mo–an. mo–he. me–ch. ca. 2–3 bo1 – do – 140 me–th*. ga–me. ch–gi–th. as in the vesicles. 2–3 bo1 the ne slope of klakkur 160 th–me*–ph. ph–ch. he–st. ca–st. 1–2 bo1 – do – 210 mo–he–ch. ch–th. th–ph. le–ch. ca–st. 1–2 bo1 – do – 260 me–th–la. an–st–la. an–th–ch. 2–3 bo1 the summit of klakkur 414 ch–th. th–ph. 2–3 bo1 the s slope of klakkur 380 as above. 2–3 bo1 – do – 300 me–th. th*–ch. th–ph–ch. he–st–la. 2–3 bo1 – do – 260 he–me–th. he–th. 2–3 bo1 the ne slope of hálgafelli 280 ce–he–me–st. ce–me–an. 2–3 bo1 – do – 300 as above. 2–3 bo1 – do – 360 as above. 2–3 bo1 – do – 380 ce–mo–he–th. ce–he–me. 2–3 bo1 – do – 400 ce–st–ch. he–th–me–st. 2–3 bo1 – do – 450 as above. 2–3 bo1 – do – 480 an–me–me*. th–ch. 2–3 bo1 the summit of hálgafelli 503 ce–he–me. st–la. 2–3 bo2 exposure at stream 0.6 km sw of norðoyri 20–80 he. me. ch. an. ap. st. cld. ca. 2–3 bo2 the w slope of høgahædd 140 an. me. ch. many empty vesicles. as above. 1–3 bo2 – do – 220 he–st. he–ch–st. th–ph. st. ca. 2–4 bo2 – do – 270 empty vesicles. – bo2 – do – 310 he–me–me*. me–an. me–ch. 2–3 bo2 – do – 320 ch–th*–ch. he–ch.th*. an. 1–2 bo2 – do – 330 me*. th’. ch. ca. 1–2 bo2 – do – 360 as above. ch. th*. ca. 1–2 bo2 – do – 440 th*. ca. op. cld. 1–2 bo2 – do – 474 me. me*. th*. ch. 1–2 bo2 – do – 510 th–th*. ca. cld. many empty vesicles. 0–1 bo2 – do – 550 as above. 0–1 bo2 the summit of høgahædd 563 as above. 0–1 bo3 large quarry between norðdepil and depil 40–50 ce–he–th. ce–he–na–th. ce–th–ch. cl–he–st–ap. cl–th– ap. cl–ph. 2–4 bo3 depilsá 150 me–me*–ch–le. me–me*–ch. he–me. 2–3 bo3 – do – 200 me*–ch, an–ch. 1–2 bo3 – do – 300 he–me*–ch. he–th*–ch. an–ph–ch. +ce. he–st. he–th–ca. 1–2 geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19153 154 profi le locality altitude (m) vesicles fractures zone bo3 depilsá 340 as above. 1–2 bo3 e slope of lokki 375 ch–gi–th’. ch–an. 0–1 bo3 – do – 400 ch–th*–ch. ch–le. 0–1 bo3 – do – 450 as above. 0–1 bo3 – do – 470 th. ch. ca. he–st–th–ca. 0–1 bo3 e slope of lokkanøv 460 ce–an–ch. ce–th–ch–th*. many empty vesicles. 0–1 bo3 – do – 580 as above. 0–1 bo3 – do – 700 as above. 0–1 bo3 the summit of lokki (trigonometric station) 754 as above. 0–1 vi1 the starting point of the profi le is the largest stream 0.5 km se of hvannasund 50 ce–me–st. ce–an–th–me. ce–me–ch. ce–he–st–me. ce–cl–st. ce–cl–me– ca. 2–3 vi1 sw slope of enni 120–132 empty vesicles. – vi1 sw slope of enni 180 mo–an. 2–3 vi1 – do – 210 an. ph–ch. an–th. me–me*. he–me– ch. he–me*–ch. he–ph– ch. he–th–me. 2–3 vi1 – do – 225 as above. 1–2 vi1 – do – 240 me*. th*. an. 1–2 vi1 – do – 270 ce–th–th*–ch. ce–ch–le. ce–an– ch. ce–ch–st. ce–he. ce–he–cld. he–st. cld. 1–2 vi1 – do – 310 ch–gi–th*. st–ch. ph–ch. th–th*. 1–2 vi1 – do – 360 ce–ch. ce–ph–ch. ce–th*. ch. 1–2 vi1 – do – 380 ch. th. le. sm. cld. 1–2 vi1 – do – 420 th*–le–ch–sm. op. sm. 0–1 vi1 – do – 550 as above. many empty vesicles. 0–1 vi1 – do – 600 ca and siderite. 0–1 vi1 the summit of enni 651 as above. 0–1 vi2 small quarry at the road hvannasund– viðareiði, 2.6 km n of hvannasund 40–80 ce–he–me. ce–he–th. ce–he– ch–sm. ce–ch–sm. ce–ph–me*. ce–me–me*. he–me–st. 2–3 vi2 w slope of tunnafjall 80–100 he–me–me*. an–th–me. an–th–ch. an–ph–ch. he–st. st–me. 2–3 vi2 – do – 150 as above. 1–2 vi2 – do – 200 as above. 1–2 vi2 – do – 225 me*. th*. ch. ph. an. he. st. ca. 1–2 vi2 – do – 250 me*. th*. ch. 1–2 vi2 – do – 300 th*. ch. 1–2 vi2 – do – 315 as above. 0–1 vi2 – do – 390 an. th’. ch. ph. me*. ph. st. 0–1 vi2 – do – 460 th*. ch. le. sm. st. sm. 0–1 vi2 – do – 520–550 empty vesicles. qz. cld. ca. 0–1 geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19154 155 profi le locality altitude (m) vesicles fractures zone vi2 the summit of tunnafjall 593 th. th*. ch. 0–1 vi2 the s slope of myrnafjall 620 no vesicles. th. th*. 0–1 vi2 the summit of myrnafjall (trigonometric station) 688 th. th*. only 20% of the vesicles are mineralised. 0–1 vi3 small quarry at the road viðareiði– hvannasund, 2.5 km south of viðareiði 80 he–th*. me–me*–ch–se. as in the vesicles. 1–2 vi3 the w and the sw slope of malinsfjall 150 he–th*. me–me*–ch. an–me*–ph. 1–2 vi3 – do – 200–220 th*. ch. le. me. me*. th. th*. an. st. 0–1 vi3 – do – 255 th*. ph. an. he–st. 1–2 vi3 the w and the sw slope of malinsfjall 310 th*–ch. 0–1 vi3 – do – 300–330 me–th*. gi–ch. 0–2 vi3 – do – 440 th*. ch. about 50% of the vesicles are empty. an–st. 0–1 vi3 – do – 540 as above. 0–1 vi3 – do – 605 ch–th–le. ca. ca–th–ca. 0–1 vi3 – do – 660 all vesicles are empty. 0–1 vi3 – do – 680 as above 0–1 vi3 – do – 710 scattered mineralisations of th and ch. 0–1 vi3 the summit of malinsfjall 750 as above. 0–1 geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19155 156 max depth min depth vesicles zone lopra-1/1a –3543 –3400 la, pr, ca, cl. ht –3400 –3200 la, mo, pr, cl. ht –3200 –3000 la, mo, pr, sm, qz. ht –3000 –2800 la, pu, qz, cl. ht –2800 –2600 (no data) –2600 –2400 la, ca, pr ,cl. ht –2400 –2200 la, pr, pu. ht –2200 –2000 th, ep, he, la, pr, wa, mo, ca, ce, sm, cl, si. 5–6 –2000 –1800 sc, th, ep, he, la, an, ca, ce, cl, si. 3–5 –1800 –1600 sc, th, st, ep, la, an, ca, ce, sm, cl, si. 3–5 –1600 –1400 sc, th, st, ep, he, la, an, ca, ce, sm, si. 4–5 –1400 –1200 me, th, st, ep, he, la, ce, sm, cl, si. 4–5 –1200 –1000 me, sc, th, ep, he, la. 3–4 –1000 –800 me, sc, th, st, ep, he, la, an, ca, cl, si. 3–4 –800 –600 th, st, ep, he, la, an, cl. 3–4 –600 –400 me, sc, th, ep, he. an. ca. 2–3 –400 –200 me, sc, th, ep, he, an, ca, cl, si. 2–3 –200 0 me, sc, th, st, he, an, mo, ca, cl, si. 2–3 vestmanna-1 –600 –575 he. he–ch. 3–4 –575 –550 he. an. st–st*. ch–sm. 3–4 –550 –525 he–ch. ap. th–ap. ch–sm. 3–4 –525 –500 he. he–ac. th–ch. st–ch. –500 –475 an–th–mt. th–sm. 3–4 –475 –450 an–th–sm. th–sm. 3–4 –450 –425 gy–th–sm. an. an–th–sm. th. sm. 3–4 –425 –400 an. he–la–ch. me–ch. th–sm. ch–sm. 3–4 –400 –375 he–la–ch. me–ap–ch. me–ch+le–sm. ch–sm. 3–4 –375 –350 me. gy–th–sm. th–sm. th*–ch. ch+le. 2–4 –350 –325 ap. th–ap. gy–la. gy–th–st. ph. th*–sm. th*–ch. 1–4 –325 –300 he. he–me. he–la–ch. mo–he–ch, th*–sm. th*–ch, ch–sm. 2–4 –300 –275 he. he–ch. me–he–st. th–ga–ch. th*–ph–ch. th*–ch. ch–sm. 2–4 –275 –250 me. me–th–ch. me–th–ph–ch.. la–me. ap. th–ch–sm. ch–sm. 2–4 –250 –225 th–gy–sm. th–th*–ch. he–th–ap. le–sm. 2–4 –225 –200 –200 –175 he. he–ch–ap–sm. me–ch, th–th*– ch. th–gy–mt. th–ap. an. ap. le–sm. 2–3 –175 –150 me–he–me*. me–he–ap. me–ap. me– he–th–le. me–th–ph–ch. th*–ch –le. na–ch. ch–sm. 2–3 –150 –125 me. me–th–ph–ch. ph–sm. me*–ch an–sm. 2–3 –125 –100 –100 –75 me. me–th–ch–sm. me+na–ap–sm. me*–ch. an–ep–ch. mo–ch. th–ch. 2–3 –75 –50 he. he–th. mo–he. me–th. na–ch. th–ch. me*–ch. ch–sm. 3–4 –50 –25 ch–sm. ep–ch. 3–4 –25 0 me. me*. me–ap. me–th. me*–ch. mo–ch. 2–3 appendix b. minerals found in vesicles within different depth intervals in the lopra-1/1a and vestmanna-1 boreholes geus bulletin no 9 7 juli.pmd 07-07-2006, 14:19156 geological survey of denmark and greenland bulletin 23, 2011, 49–52 49 any major shaking of the earth can be recorded on a seismograph regardless of the nature of the source. earthquakes and large explosions generate waves with similar frequency content. this fact has been used for decades to construct systems to monitor detonations of underground nuclear explosions. the quality of the monitoring system has increased significantly in recent years, and we demonstrate here that the data are useful in danish earthquake research. one important difference between explosions and earthquakes is the depth of the source, most earthquakes occurring at much larger depths than explosions. thus the depth determination is important in both fields. however, accurate depth determination of earthquake hypocentres is a challenge even when it comes to large well-recorded earthquakes. the uncertainty of the calculated depth of a danish earthquake is of the same magnitude as the depth itself when using standard location techniques. a technique utilising crustal phases recorded at large distances has been introduced at the geological survey of denmark and greenland to improve the determination of the hypocentre depths. only the largest earthquakes in denmark and its immediate surroundings produce sufficiently strong signals to be recorded at teleseismic distances, i.e. larger than 3000 km. the signals are discernable at low-noise seismic array stations as far away as north america and africa. some of these stations are operated by the united nations comprehensive nuclear-test-ban treaty organisation (ctbto). data from many of these stations are available for scientific purposes. we demonstrate here how data from canada and niger can significantly improve the depth estimates for two earthquakes, one in skåne, sweden in 2008 and the other in the danish part of the north sea in 2010 (figs 1, 2). ctbto and ims the comprehensive nuclear-test-ban treaty was adopted by the general assembly of the united nations (un) in september 1996. the treaty bans all nuclear explosions on earth whether for military or for peaceful purposes. denmark signed the treaty in 1996, and it was ratified by the parliament of denmark in 1998. the treaty will enter into force once it has been signed and ratified by all nuclear powers of the world. however, the ratification is still pending in several key countries such as china, pakistan, india and united states of america. comprehensive nuclear-test-ban treaty – a peace-keeping initiative with scientific impact tine b. larsen, peter h. voss, trine dahl-jensen and søren gregersen 10°e 14°e 55°n 57°n 100 km sweden denmark germany skåne earthquake north sea earthquake yka norsar tord spits hfs fines arces fig. 1. map of the north atlantic region showing the locations of seismic array stations used by geus to locate earthquakes in denmark and greenland. yka, canada. spits, norsar and arces, norway. fines, finland. hfs, sweden. tord, niger. fig. 2. teleseismic array data have been used to locate earthquakes and measure their depths in skåne on 16 december 2008 and in the north sea on 19 february 2010. © geus, 2011. geological survey of denmark and greenland bulletin 23, 49–52. open access: www.geus.dk/publications/bull 5050 significant resources are being allocated to the development of a large international monitoring system (ims), so that all technical systems as well as procedures, manuals and other agreements are in place once the treaty enters into force. the ims is a un-controlled monitoring system based on four technologies: seismology, radio nuclides, hydroacoustics and infrasound. the system consists of a worldwide network of high-quality monitoring stations, supplementing the existing national networks of detectors, and it has improved the detection threshold for explosions as well as earthquakes in large parts of the earth. data from the stations are transmitted via secure satellite links to a un data centre in vienna. all the data are processed at the centre in vienna. the data centre is completely neutral and is not permitted to judge if an event is natural, such as an earthquake or a volcanic eruption, or caused by a man-made explosion. instead the data centre makes raw data as well as processed data available to the individual countries, and then it is up to the national authority in each country to decide whether an event is suspicious or not. should an event be deemed suspicious by a country, any country has the right to request further processing by the ims and ultimately request an on-site inspection at the location of the suspicious event. in recent years, nuclear test explosions have been easy to identify, as the involved nations have openly announced the tests and provided information about location and time of the explosions (pakistan 1998, india 1998, north korea 2006 and 2009). geus and ctbto in denmark, geus houses both the national data centre and is also the national authority for the ctbto. geus is responsible for running two monitoring stations in the ims network. one is a seismograph in kangerlussuaq, west greenland; it is part of the auxiliary (secondary) seismic network. the other is an infrasound station near qaanaaq, north-west greenland, that is a primary ims station. the danish meteorological institute takes part in the daily maintenance of the infrasound station. together with diplomats from the danish embassy in vienna, geus is also involved in the preparatory work on operational manuals and procedures for waveform processing that is carried out at the un centre. this involves discussions on how to tune the detection system to automatically filter out as many earthquake signals as possible, so that the system triggers on explosions only. this part of the system still needs significant improvements. in order to be as familiar as possible with the data from the ims before the treaty enters into force, seismologists at geus make experiments with the incoming data for other purposes. recently we have found that the ims raw seismological data are useful for determining accurate depths of danish earthquakes. the ims seismographs are of a very high quality and it is possible to identify signals from relatively weak earthquakes at distant stations. earthquake depth analysed using ctbto array data for several decades we have supplemented with data from seismic stations in the countries around denmark and greenland to help detect and localise earthquakes. seismic array stations produce data of particularly high quality. a seismic array station consists of a large number of sensors installed in a small area. the ability to detect an earthquake decreases when the strength of the seismic noise increases relative to the strength of the seismic signal. seismic noise can be reduced significantly in earthquake data recorded by an array by summing the signals recorded at the sensors, thus enabling the detection of smaller events than is possible on a station with just one sensor. the closest seismic arrays that contribute to the monitoring of earthquakes in denmark and greenland are the norsar and arces arrays in norway, the hfs array in sweden, the fines array in finland and the spits array at svalbard (fig. 1). previously, array data have been used to locate earthquakes in denmark and greenland from observations of p and s phase travel times only. for the two earthquakes in skåne on 16 december 2008 and in the north sea on 19 february 2010, however, a new technique was used to improve the estimate of the depth of the earthquakes. this fig. 3. the slightly different paths taken by the pand the pp(or sp-) waves can be used to calculate the depth of an earthquake. the p-wave (yellow) travels directly from the hypocentre to the seismograph, whereas ppand sp-waves (green) travel from the hypocentre to the surface near the epicentre and from there reflected to the seismograph. the paths of the ppand sp-waves differ slightly due to the crustal structures near the earthquake. epicentre seismograph hypocentre p-wave pp-wave 50–70° 51 technique uses teleseismic observations of travel-time differences between the p phase and the pp and/or sp phases to calculate the depth of the earthquake. the pp and sp phases are reflections at the earth’s surface of the shaking from the earthquake hypocentre. these phases are recorded slightly after the p wave arrival (fig. 3). the technique requires measurements with a high signal-to-noise ratio and good knowledge of the geological structures near the epicentre. the principles of the teleseismic depth-determination technique are very simple. at teleseismic distances the difference in travel length is negligible for the p-wave travelling directly from the hypocentre to the seismograph and the part of the pp-wave travelling from the epicentre to the seismograph (fig. 3). the difference in travel time between the phases is assumed to be caused by the pp-wave travelling almost vertically through the crust from the hypocentre to the surface. if the velocity structure in the crust below the epicentre is well known, the travel time difference between p and pp and/or sp can be converted to a depth. an accurate crustal velocity model is therefore critical for the analysis since errors in the crustal model will give a wrong determination of the depth. a teleseismic recording is needed as the paths of the pp and sp phases must be near vertical at the source. however, the distance between the earthquake and the seismic array must be sufficiently short for the phase not to be effected by the core–mantle boundary. we find that a distance of around 50 to 70 degrees is optimal for this technique. a significant number of the ims stations are seismic arrays. the stations are installed at locations with low ambient noise, in order to record data of high quality. in our analysis we use data from the yellowknife array (yka) in canada and the toridu array (tord) in niger (fig. 1). the yellowknife array consists of 19 short-period sensors and 4 broadband sensors. the sensors are installed in a cross with an aperture of 25 km. the array was installed in 1962 with the main purpose of monitoring underground nuclear explosions (source: natural resources canada). the toridu array is a modern array constructed specifically for the ctbto (estabrook et al. 2009). it consists of 16 broadband sensors deployed in three concentric rings with a central node. the sensors are not radially aligned, as this layout leads to the largest noise reduction (e.g., schweitzer et al. 2002). as described above, the energy release of the majority of the earthquakes in denmark and greenland is too low to generate clear signals even at the best ims stations. the two earthquakes, skåne, 16 december 2008 and the north sea, 19 february 2010, measuring 4.8 and 4.7 on the richter scale, respectively, are our best candidates for this technique (fig. 2). the geological structures are well mapped in denmark where we have a good knowledge of p-wave velocity (thybo 2001). for the s-wave velocity we have used vp = 1.73 × vs to estimate the earthquake depth. using this technique in greenland will be less reliable in most areas, because the crustal structures are not as well mapped. we have analysed the measurements of the skåne earthquake on the yka array (fig. 4). the measurements show a good signal-to-noise ratio and we find a difference in the p and sp travel times of c. 4 sec., equivalent to an earthquake depth of 9 km. the previously calculated depth of this earthquake was 18.1 ± 5.2 km using the standard location method (snsn 2010), and the result from moment-tensor inversion was 8 km (regel 2010). the standard location method is based on an approach that searches for the hypocentre that gives the best fit to measured travel times of pand s-waves within c. 1000 km of the earthquake, using a 1d earth model. the larger depth obtained by snsn might be due to the velocity model, which is not well calibrated for the skåne 19 feb. 2010 21:16:50.0 utc 55 05 15 25 toa2 tob1 tob2 tob3 tob4 tob5 toc1 toc2 toc3 toc4 toc5 toc6 toc7 16 dec. 2008 05:29:31.9 utc sec.36 40 44 48 yka data, vertical component, unfiltered ykb0 ykb1 ykb3 ykb6 ykb7 ykb8 ykb9 ykr1 ykr3 ykr4 ykr5 ykr7 ykw3 tord data, vertical component bandpass filtered 1–5 hz a b plot start time: fig 4. a: the p-wave train from the skåne earthquake on the yellowknife array (yka). yellow is the direct p phase and green is the sp phase. b: the p-wave train from the north sea earthquake on the toridu array (tord). yellow is the direct p phase and green is the pp phase. the high number of sensors improves the possibility of identifying different phases. 5252 area (b. lund, personal communication 2010) and the use of measurements far from the epicentre. the north sea earthquake measurements on the yka have a low signal-to-noise ratio on many of the sensors, but on the best five sensors we observe a signal c. 15 sec. after the p phase. interpreting this as the sp phase yields a depth of 35.2 km. the high noise level in the yka data carries a risk of misinterpretation of the data. we therefore supplement with data from the tord array (fig. 4), to verify the depth of this earthquake. from the measurements of the tord array we find a difference in the p and pp travel time of c. 11 sec., corresponding to an earthquake depth of 38.5 km. the previously calculated depth was 38.7 ± 10.3 km using the standard location method. the depth uncertainty is larger than that of the skåne earthquake due the larger distance of the nearest seismometer. the depths of the two earthquakes are comparable with the depths of previous earthquakes in these areas (gregersen et al. 1999). the higher noise level of the north sea event (fig. 4) is a source of error in the analysis. the higher noise level could be due to an energy radiation pattern of the earthquake that is low in the direction of tord or a different frequency content of the released shaking. concluding remarks teleseismic array data have produced consistent depth estimates for two recent earthquakes in the danish area. this raises the possibility that we might find other earthquakes suitable for this technique in the geus database, especially from greenland. the seismological involvement in detection and discrimination of nuclear explosions has spurred significant nordic collaboration since the 1960s, and a yearly nordic seismological meeting is held. geus has made a special contribution together with the uk foreign and commonwealth office, norsar, the swedish national defence research establishment and the university of helsinki to improve data exchange between the international monitoring system and the international seismological centre. this has resulted in the development of a collection of interactive seismological tools for merging and manipulating the two largest and most complete seismological databases (gaspa et al. 2010). in the future this will hopefully help improve and ease the scientific use of the international monitoring system data. references comprehensive nuclear test-ban treaty (www.ctbto.org). estabrook, c., bergsson, b., soumana, s., boureima, o. & moumouni, m. 2009: results from ims seismic array in niger. poster presented at the egu meeting in vienna, april 2009. gaspa, o., bondar, i., harris, j. & storchak, d. 2010: the ctbto link to the isc database. the 41st nordic seminar on detection seismology, århus, 6–8 october, 2010. program with abstracts, 14 only. gregersen, s., hjelme, j. & hjortenberg, e. 1998: earthquakes in denrnark. bulletin of the geological society of denmark 44, 115–127. regel, j. 2010: moment tensor of the 16 dec[ember] 2008 earthquake in skåne, sweden. the 41st nordic seminar on detection seismology, århus, 6–8 october 2010. program with abstracts, 39 only. schweitzer, j., fyen, j., mykkeltveit, s. & t. kværna, t. 2010: seismic arrays. in: new manual of seismological observatory practice, chapter 9. doi: 10.2312/gfz.nmsop_rl_ch9. snsn 2010: swedish national seismic network (http://snsn.geofys. uu.se/). thybo, h. 2001: crustal structure along the egt profile across the tornquist fan interpreted from seismic, gravity and magnetic data. tectonophysics 334, 155–190. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: tbl@geus.dk geological survey of denmark and greenland bulletin 20, 2010, 19–22 19 concerns about climate change have led to increased interest in geothermal energy as one way of reducing the consumption of fossil fuels and thus limit co2 emissions. use of geothermal energy is based on well-established technologies, a high degree of security of supply, and little visual or noise inconvenience. more than one hundred plants have been established in europe. there is a large potential for using geothermal energy from the danish subsurface, as first pointed out by balling (1976). geothermal energy is highly suitable for district heating systems and is expected to cover a large part of the demand for district heating in the future. two danish geothermal plants, the thisted plant in northern jylland and the margretheholm demonstration plant near copenhagen (fig. 1), have shown that it is possible to produce large amounts of warm water for district heating. only 5–10% of the total energy output from the plant is used to extract the heat from the subsurface by pumping warm formation water to the surface and returning it to the subsurface in a closed system. the plants use absorption warmth pumps, which need steam and hence give rise to consumption of (fossil) fuel. both danish plants have two wells, a production well and an injection well in which the cooled formation water is returned to the geological reservoir at about 1 km away from the production point, in order to avoid mixing of warm and cold water (fig. 2). geothermal energy can also be used for electricity production, but danish subsurface temperatures are currently not believed to be sufficiently high to produce electricity directly. because geothermal energy is expected to play an increasingly important role in the energy strategy of denmark, the identifying potential geothermal reservoirs in denmark anders mathiesen, lars henrik nielsen and torben bidstrup frederikshavn reservoir haldager reservoir gassum reservoir skagerrak reservoir bunter reservoir 12°e major fault well geothermal plant structural high gassum res. too deep bunter res. too deep reservoirs too shallow r i n g k ø b i n g – f y n h i g h danish basin sweden ø resund n o r t h g e r m a n b a s i n sorgenfrei–tornquist zone skagerrak–kattegat platform 56°n 55°n 50 km 8°e 10°e 57°n thisted viborg jylland sjælland sønderborg margretheholmfig. 1. the distribution of potential sandstone reservoirs in denmark with depths in the 800–3000 m interval and thicknesses above 25 m. dark grey and black areas indicate where the reservoirs are too deep (gassum in northern jylland; bunter in western jylland – both located in the central parts of the danish basin). light grey areas indicate where reservoirs are absent (ringkøbing–fyn high) or too shallow (less than c. 800 m; northernmost jylland). hatched areas indicate two or more reservoirs with geothermal potential. green areas: not mapped. © geus, 2010. geological survey of denmark and greenland bulletin 20, 19–22. open access: www.geus.dk/publications/bull 2020 geological survey of denmark and greenland (geus) and the danish energy agency have conducted a regional study to update the assessment of the geothermal potential in denmark (mathiesen et al. 2009). based on existing well, seismic and temperature data and the detailed knowledge of the subsurface stratigraphy gathered by geus over many years, the assessment has documented a huge geothermal potential in many parts of denmark. the focus of the study was to evaluate (1) the potential of geothermal energy in denmark and (2) if it can contribute significantly to the danish strategy for a safe, sustainable and reliable supply of energy. the specific potential in local areas was not evaluated in detail; however, a well-defined and stepwise procedure to develop local geothermal prospects by integrating existing and new data is suggested. potential reservoirs and areas of interest the danish subsurface can be divided into five major structural parts: the north german basin, the ringkøbing–fyn high, the danish basin, the sorgenfrei–tornquist zone and the skagerrak–kattegat platform (fig. 1). these structural divisions exert a decisive influence on the geothermal prospectivity of the danish subsurface, as they essentially determine the distribution, thicknesses, facies types and burial depths of the potential reservoirs (nielsen 2003; nielsen et al. 2004). the 1–10 km thick mesozoic succession has been the target for hydrocarbon exploration since 1935 and is thus relatively well known from about 60 deep wells and seismic data acquired over many years, although with a highly variable data coverage and quality. the data show that the most promising geothermal reservoirs occur within the triassic – lower cretaceous succession in the danish basin and the north german basin, separated by the ringkøbing–fyn high which shows a lower potential. based on regional geological studies, four main stratigraphical units with a regional geothermal potential have been identified (nielsen et al. 2004). within these four main units, we have defined five geothermal reservoirs on the basis of their stratigraphical and spatial extent. each reservoir comprises a large number of sandstone layers that are potential aquifers. the new assessment shows that large areas in both basins have a good geothermal potential, as they contain several porous, water-bearing sandstone reservoirs in the economic interval 800–3000 m below the surface with formation temperatures of 25–90°c (mathiesen et al. 2009). the principal challenge for successful geothermal exploration is to assess whether good reservoir properties are present in terms of sufficient layer thickness, porosity, permeability, temperature and formation water geochemistry. these geological factors are used to evaluate whether the thermal energy of the formation water can be exploited economically and used for heating purposes. by combining knowledge of the geographical distribution of the stratigraphical units with the potential reservoirs, their mapped burial depths and estimates of where the cumulative thickness of the reservoir sandstones exceeds c. 25 m at depths of 800–3000 m, a useful indication of the regional geothermal potential is provided. temperature and salinity of the formation water in the potential reservoirs increase with increasing depth. based on data from a number of wells a general temperature–depth relation has been established, showing a gradient of 25–30°c/km. the salinity of the formation water shows a general increase of about 10%/km burial depth, but large variations are found. porosity and permeability decrease with increasing depth due to mechanical compaction and formation of diagenetic minerals that reduce the pore volume (porosity) and the connections between the pores (permeability). several of these properties are directly related to the petrography of the sediment source areas and the grain size distribution of the material supplied to the basins. thus the various depositional processes during the formation of the reservoirs and their subsequent burial depths determine their qualities as geothermal reservoirs. however, the mutual dependency of the various factors and processes is not fully understood, which weakens the predictive strength and reliability of the geological models currently in use to identify areas of interest. figure 1 shows regions where sandstone-rich reservoirs are expected to have geothermal potential by combining geographical distribution and burial depth information of the five reservoirs. the map also indicates where further detailed research and investigations are required if geothermal enerreservoir consumergeothermal plant with heat exchanger fig. 2. schematic diagram illustrating the geothermal concept used in denmark. in the geothermal plant a production well pumps up warm water (red line) from the underground reservoir, and an injection well (blue line) returns the cooled water to the reservoir. 21 gy production is to be further developed (mathiesen et al. 2009). the central part of the danish basin is promising, whereas the potential of areas located around the 800 m and 3000 m cut-off limits is highly uncertain. areas located along the ringkøbing–fyn high are considered less prospective. the need for further development interest in the use of geothermal energy has increased over the past five years. so far, the thisted and margretheholm plants are the only working geothermal plants in denmark. however, geus has carried out evaluations of the geothermal potential in several local areas in denmark with positive conclusions. it is expected that new danish geothermal plants will be established in sønderborg and viborg within the next few years. the existing thisted plant has produced heat from the gassum formation (the gassum reservoir) for almost 20 years without notable production or injection problems, and the newly established margretheholm plant produces heat from the bunter sandstone formation (the bunter reservoir). during planning of the margretheholm plant new seismic data were acquired in 2001, and geus carried out a geological evaluation of the geothermal potential. the study used old and new seismic data to map the distribution, the tops and the lateral variations of the reservoirs, as well as faults. faults may reduce the lateral continuity of a reservoir (fig. 3). from this study it was concluded that the bunter reservoir is found south of the ringkøbing–fyn high, on parts of the high and in the danish basin, and grades into the skagerrak reservoir towards the north-eastern basin margin (fig. 2; mathiesen et al. 2009). the bunter reservoir is dominated by fine-grained sandstones that were deposited in arid continental environments with fluvial channels, aeolian dunes and with some marine influence. in southern sweden, analyses of existing cuttings from old wells and data from new wells documented the presence of a loose, mediumto coarse-grained, quartzitic sandstone composed of sub-rounded quartz and feldspar grains without overgrowth of minerals or other signs of corrosion and cementation. data from the swedish wells combined with log correlation to danish wells indicated the presence of a c. 50 m thick, loose, conglomeratic sandstone unit in the basal part of the bunter reservoir, and this unit was predicted to show high porosity and low degree of cementation at margretheholm. in 2002 and 2003 the margretheholm-1 and -2 wells were drilled to about 2700 m depth and confirmed the presence of several sandstone-rich aquifers in both the gassum and bunter reservoirs. studies of the new wells and results from log correlation with other wells in the copenhagen region, øresund and southern sweden confirmed the previous geological model of the bunter reservoir and strengthened the indications of a large geothermal potential in the copenhagen area (figs 3, 4). no cores were taken from the new wells, but both wells encountered a promising aquifer in the bunter reservoir with satisfying test results, and in 2006 the marsw ne mah-1 geothermal well t w o -w ay t ra ve l ti m e (m se c) 1500 1000 500 base upper cretaceous top gassum top bunter top basement 2 km amager fault fig. 3. sw–ne-trending seismic profile south-west from margretheholm. the profile illustrates the importance of mapping faults, which can reduce the lateral reservoir continuity. for location see fig. 4. mah-1/-2 12°20´e 12°40´e 55°30´n 55°45'n55°45'n55°45´n 10 km fig. 3 1117–1200 1201–1400 1401–1600 1601–1800 1801–2000 2001–2200 2201–2400 2401–2600 2601–2800 2801–3000 3001–3200 3201–3400 seismic line depth to top bunter reservoir fig. 4. depth to the top of the bunter reservoir, illustrating southwards shallowing of the reservoir top and its partition by faults (grey tone). red dot: margretheholm and the wells mah-1 and mah-2. 2222 data, which vary highly in density and quality, but also on our understanding of the geological processes that led to the formation of the geothermal reservoirs. a large, newly funded, multi-disciplinary research project carried out jointly by geus, the university of aarhus, the geological survey of sweden, deutsches geoforschungs zentrum and dong energy aims to assess these challenges in detail in order to encourage the utilisation of geothermal energy by investigating the critical elements in the exploitation of the geothermal resources and integrating the results on a gis platform. the project will use existing and new geological and geophysical data and methods. by conducting detailed studies of relevant, sandstone-rich reservoirs and the diagenetic processes that affected the reservoirs since deposition, the project will provide new data leading to improved understanding of the spatial variations in the physical characteristics and quality of the reservoirs. new seismic data will be acquired, and interpretations of these data and old data will increase our knowledge about lateral reservoir variations and will form the basis of a new consistent 3-d subsurface temperature model. thermal and fluid flow modelling using various schemes of combined production and re-injection wells can provide more detailed information on the thermal resources that can be extracted from the reservoirs and their thermal life-time, and it will be possible to quantify the additional amount of heat that can be extracted from less prospective layers both within and adjacent to the geothermal reservoirs. traditionally this is not included in estimates of resources, but it will provide higher and more realistic values. references balling, n. 1976: geothermal models of the crust and uppermost mantle of the fennoscandian shield in south norway and the danish embayment. journal of geophysics 42, 237–256. mathiesen, a., kristensen, k., bidstrup, t. & nielsen, l.h. 2009: vurdering af det geotermiske potentiale i danmark. danmarks og grønlands geologiske undersøgelse rapport 2009/59, 30 pp. nielsen, l.h. 2003: late triassic – jurassic development of the danish basin and the fennoscandian border zone, southern scandinavia. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 459–526. nielsen, l.h., mathiesen, a. & bidstrup, t. 2004: geothermal energy in denmark. geological survey of denmark and greenland bulletin 4, 17–20. gretheholm geothermal plant was opened based on water with a temperature of 73°c. despite the encouraging results, a more detailed study of the new margretheholm wells is required, including correlation with other wells in the danish area and with the wellknown bunter sandstein in the german part of the north german basin to be able to estimate the potential of other sandstone-rich layers in the bunter reservoir, both in the copenhagen area and in other places in the southern part of denmark. further studies are needed because it has been decided that the copenhagen area should be co2-neutral by 2025. to achieve this goal, the margretheholm plant is expected to play a major role in the use of geothermal energy for district heating. this will require an expansion of the plant over the next five years. therefore a more detailed evaluation is necessary, both to minimise the prospected risk and to enhance the optimal use of the subsurface below copenhagen for geothermal energy. future potential – towards a new resource assessment a more environmentally friendly denmark has to include geothermal energy as a significant part of the energy supply. it is estimated that the geothermal resources in denmark can last several hundred years with the present heat consumption, and only a small fraction of this potential is utilised by the two existing geothermal power plants (mathiesen et al. 2009). a major challenge within geothermal prospecting is to find suitable and sustainable reservoirs (small number of faults and lateral facies changes) and sufficient flow capacity (thickness, porosity and permeability) of warm water. permeability is critical but difficult to predict since large variations occur depending on the depositional facies, provenance, mineralogical composition, burial history and position in the basin. one of the barriers for a significant increase in the exploitation of the large geothermal resources in denmark is the geological uncertainty in the exploration phase. this uncertainty is related to the possibility of making accurate and reliable predictions about the presence of high-quality reservoirs with high lateral continuity below the urban areas, where infrastructure and consumers are located. precise predictions depend not only on existing well and seismic authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: anm@geus.dk 1 geological survey of denmark and greenland bulletin 32 • 2014 a catalogue of danian gastropods from the baunekule facies, faxe formation, denmark bodil wesenberg lauridsen & kai ingemann schnetler geological survey of denmark and greenland danish ministry of climate, energy and building 22 geological survey of denmark and greenland bulletin 32 keywords cold-water coral mound ecosystem, danian, gastropods, new taxon, palaeoecology, systematics cover illustration the faxe limestone quarry on the stevns peninsula, eastern denmark, has been worked intermittently since the middle ages. intensive collecting by quarrymen as well as amateur and professional palaeontologists has resulted in rich fossil collections providing a detailed insight into the diversity of life on and around cold-water coral mounds in danian times (about 63 million years ago). the quarry is viewed towards the north-east from close to the recently-established geomuseum faxe. frontispiece: facing page original drawings of the danian gastropods from the faxe quarry were undertaken in the 1960s under the direction of professor a. rosenkrantz; these are now stored together with the specimens in the geological museum of the university of copenhagen. this page (rkz 64) shows four drawings of the tiny gastropod cerithiscala tricincta (ravn 1933), only a few millimetres in height, which is illustrated in this catalogue in fig. 73. chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, department of geoscience, aarhus university; minik rosing, geological museum, university of copenhagen; finn surlyk, department of geography and geology, university of copenhagen scientific editor of this volume: jon ineson editorial secretary: jane holst referees: j. alistair crame (uk) and thomas hansen (no) illustrations: jette halskov with contributions from adam a. garde digital photographic work: benny schark graphic production: kristian a. rasmussen printers: rosendahls-schultz grafisk a/s, albertslund, denmark manuscript received: 5 may 2014 final version approved: 1 november 2014 printed: 30 december 2014 issn 1604-8156 isbn 978-87-7871-391-9 citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 32, 117 pp. errata (as of 19 february 2015) a number of figured specimens in the publication have been labeled with an incorrect institutional affiliation. these have been corrected in this electronic version, but not in the printed version; an errata supplement has been inserted in the printed bulletins. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk and at www.geus.dk/publications/bull © de nationale geologiske undersøgelser for danmark og grønland (geus), 2014 for the full text of the geus copyright clause, please refer to www.geus.dk/publications/bull 3 44 contents abstract . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 15 introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17 material . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17 previous work on gastropods from the baunekule facies . . . . . . . . . . . . . . . . . . . . . . . 18 geological setting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19 notes on the locality and the baunekule facies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 20 remarks on the fauna . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23 preservation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23 number of taxa . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23 stratigraphic ranges of the genera . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23 comparison with other paleocene gastropod faunas . . . . . . . . . . . . . . . . . . . . . . . . . 23 palaeoecology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 28 gastropods on modern cold-water coral mounds . . . . . . . . . . . . . . . . . . . . . . . . . . . . 31 concluding remarks and future studies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 32 acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 32 systematic palaeontology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 35 abbreviations and classification . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 35 phylum mollusca linnaeus 1758 class gastropoda cuvier 1795 clade vetigastropoda superfamily fissurellioidea fleming 1822 family fissurellidae fleming 1822 subfamily fissurellinae fleming 1822 genus rimula lowe 1852 rimula? sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 36 subfamily emarginulinae children 1834 genus emarginula lamarck 1801 emarginula coralliora lundgren 1867 . . . . . . . . . . . . . . . . . . . . . . . . . . 36 superfamily pleuroromarioidea swainson 1840 family pleurotomariidae swainson 1840 genus leptomaria eudes-deslongchamps 1864 leptomaria niloticiformis (von schlotheim 1820) . . . . . . . . . . . . . . . . . . 36 leptomaria sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 37 superfamily haliotoidea rafinesque 1815 family temnotropidae cox 1960 genus temnotropis laube 1870 temnotropis sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 37 superfamily porcellioidea koken 1895 family porcellidae koken 1895 genus faxetrochus schnetler & lozouet 2012 faxetrochus problematicus schnetler & lozouet 2012 . . . . . . . . . . . . . . . 37 superfamily scissurellioidea gray 1847a family scissurellidae gray 1847a subfamily scissurellinae gray 1847a 5 genus scissurella d'orbigny 1824 scissurella annulata ravn 1933 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 38 scissurella (s. l.) aliceae schnetler, lozouet & pacaud 2001 . . . . . . . . . . . . 38 scissurella ravni (schnetler, lozouet & pacaud 2001) . . . . . . . . . . . . . . . 39 family anatomidae mclean 1989 genus anatoma woodward 1859 anatoma hedegaardi (bandel 1998) . . . . . . . . . . . . . . . . . . . . . . . . . . . 39 anatoma rosenkrantzi schnetler, lozouet & pacaud 2001 . . . . . . . . . . . . 40 superfamily trochoidea rafinesque 1815 family trochidae rafinesque 1815 subfamily monodontinae gray 1857 genus osilinus philippi 1847 osilinus carinatus (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 40 genus jujubinus monterosato 1884 jujubinus sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 40 subfamily stomatellinae gray 1840 genus stomatella lamarck 1816 stomatella sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 41 family turbinidae rafinesque 1815 subfamily skeneinae clark 1851 genus leucorhynchia crosse 1867 leucorhynchia marginata ravn 1933 . . . . . . . . . . . . . . . . . . . . . . . . . . . 41 genus skeneoides warén 1992 skeneoides sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 42 family ataphridae cossmann 1915 genus ataphrus gabb 1869 ataphrus sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 42 subfamily margaritinae thiele 1924 genus margarites gray 1847a margarites bruennichi (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . . . . 43 superfamily angarioidea williams et al. 2008 family angariidae gray 1857 subfamily angariinae gray 1857 genus angaria röding 1798 angaria depressa (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 43 superfamily phasianelloidea williams et al. 2008 family colloniidae cossmann in cossmann & peyrot 1917 subfamily colloninae cossmann in cossmann & peyrot 1917 genus collonia j.e. gray 1850 collonia (circulopsis) pusilla ravn 1933 . . . . . . . . . . . . . . . . . . . . . . . . 43 genus vexinia cossmann 1918 vexinia sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 44 superfamily eucycloidea koken 1897 family eucycloscalidae gründel 2007 genus eucycloscala cossmann 1895 eucycloscala ultima ravn 1933 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 44 66 superfamily seguenzioidea verril 1884 family chilodontidae wenz 1938 genus danilia brusina 1865 danilia faxensis (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 44 danilia quadricordata (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . . . 45 danilia fenestrata (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 45 subfamily chilodontinae wenz 1938. chilodontinae gen. et sp. indet. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 45 clade caenogastropoda cox 1960 clade sorbeoconcha ponder & lindbergh 1997 superfamily cerithioidea fleming 1822 family cerithiidae fleming 1822 genus ataxocerithium tate 1894 ataxocerithium faxensis (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . . . 46 ataxocerithium sp. 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 46 ataxocerithium sp. 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 46 genus bittium leach in gray 1847b bittium cf. transenna (bayan 1873) . . . . . . . . . . . . . . . . . . . . . . . . . . . . 47 family metacerithiidae cossmann 1906 genus metacerithium cossmann 1906 metacerithium? sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 47 superfamily campaniloidea douvillé 1904 family campanilidae douvillé 1904 genus campanile bayle in fischer 1864 campanile? pseudotelescopium (ravn 1902b) . . . . . . . . . . . . . . . . . . . . . 48 campanile? subglabra (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . . . . 48 campanile? sp. 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 48 campanile? sp. 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 49 family trypanaxidae gougerot & le renard 1987 genus trypanaxis cossmann 1889 trypanaxis faxensis ravn 1933 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 49 family turritellidae lovén 1847 genus mesalia gray 1847a mesalia? sp. 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 50 mesalia? sp. 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 50 mesalia? sp. 3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 51 family siliquariidae anton 1838 genus tenagodus guettard 1770 tenagodus ornatus (lundgren 1867) . . . . . . . . . . . . . . . . . . . . . . . . . . . 51 tenagodus sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 51 clade hypsogastropoda ponder & lindberg 1997 clade littorinimorpha golikov & starobogatov 1975 superfamily capuloidea fleming 1822 family capulidae fleming 1822 7 genus trichotropis broderip & sowerby 1829 trichotropis? sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 52 capulidae gen. et sp. indet. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 52 superfamily cypraeoidea rafinesque 1815 family cypraeidae rafinesque 1815 subfamily gisortiinae schilder 1927 genus palaeocypraea schilder 1928 palaeocypraea spirata (von schlotheim 1820) . . . . . . . . . . . . . . . . . . . . . 52 palaeocypraea sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 53 genus bernaya jousseaume 1884 subgenus bernaya (protocypraea) schilder 1927 bernaya (protocypraea) globuliformis (ravn 1902b) . . . . . . . . . . . . . . . . . 53 family eocypraeidae schilder 1924 genus eocypraea cossmann 1903 eocypraea danica (schilder 1928) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 54 superfamily littorinoidea children 1834 family littorinidae children 1834 subfamily littorininae children 1834 genus littoraria gray 1833 subgenus littoraria (littorinopsis) mörch 1876 littoraria (littorinopsis) faxensis (ravn 1933) . . . . . . . . . . . . . . . . . . . . 54 family pickworthiidae iredale 1917 subfamily pickworthiinae iredale 1917 genus sansonia jousseaume 1892 sansonia hedegaardi bandel & kowalke 1997 . . . . . . . . . . . . . . . . . . . . 55 sansonia sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 55 genus mareleptopoma moolenbeek & faber 1984 mareleptopoma? sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 55 genus urceolabrum wade 1916 urceolabrum sp. 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 55 urceolabrum sp. 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 56 subfamily sherboniinae iredale 1917 genus faxia ravn 1933 faxia macrostoma ravn 1933 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 56 superfamily naticoidea guilding 1834 family naticidae guilding 1834 naticidae gen. et sp. indet. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 56 superfamily rissooidea gray 1847a family rissoidae gray 1847a subfamily rissoininae gray 1847a genus zebina h. & a. adams 1854 zebina sp. 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57 zebina sp. 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57 genus pseudotaphrus cossmann 1888 pseudotaphrus sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57 rissoininae gen. et sp. indet. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 58 88 family caecidae gray 1850 genus caecum fleming 1817 caecum sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 58 family tornidae sacco 1896 (1884) genus circulus jeffreys 1865 circulus sp. 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 58 circulus sp. 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 58 subfamily teinostomatinae cossmann in cossmann & peyrot 1917 genus teinostoma h. & a. adams 1853 teinostoma glaberrimum ravn 1933 . . . . . . . . . . . . . . . . . . . . . . . . . . . 59 genus sigaretornus iredale 1936 sigaretornus sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 59 superfamily tonnoidea suter 1913 (1825) family ranellidae gray 1854 genus ranella lamarck 1816 ranella faxensis ravn 1933 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 60 subfamily cymatiinae iredale 1913 (1854) genus sassia bellardi 1873 sassia faxense (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 60 sassia sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 60 genus cymatium röding 1798 subgenus cymatium (monoplex) perry 1811 cymatium (monoplex) subglabrum (ravn 1902b) . . . . . . . . . . . . . . . . . . 61 superfamily truncatelloidea gray 1840 family elachisinidae ponder 1985  genus laeviphitus van aartsen, bogi & giusti 1989 laeviphitus sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 61 superfamily vanikoroidea gray 1840 family hipponicidae troschel 1861 genus hipponix defrance 1819 hipponix sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 62 genus eoatlanta cossmann 1889 eoatlanta ravni schnetler 2013 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 62 superfamily velutinoidea gray 1840 family triviidae troschel 1863 subfamily eratoinae gill 1871 tribe johnstrupiini schilder 1939 genus johnstrupia ravn 1933 johnstrupia faxensis ravn 1933 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 63 informal group ptenoglossa superfamily epitonioidea berry 1910 family epitoniidae berry 1910 genus cerithiscala de boury 1887 cerithiscala tricincta (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 64 cerithiscala sp. 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 64 9 cerithiscala sp. 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 64 genus opalia h. & a. adams 1853 opalia sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 65 genus acrilla h. adams 1860 acrilla elegans (ravn 1902b) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 65 acrilla sp. 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 66 acrilla sp. 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 66 acrilla sp. 3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 67 acrilla? sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 67 genus cirsotrema mörch 1852 cirsotrema sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 67 superfamily eulimoidea philippi 1853 family eulimoidae philippi 1853 genus eulima risso 1826 subgenus eulima (polygireulima) cossmann 1894 eulima (polygireulima) danica ravn 1933 . . . . . . . . . . . . . . . . . . . . . . . 68 eulima (polygireulima) sp. 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 68 eulima (polygireulima) sp. 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 69 eulima (polygireulima) sp. 3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 69 eulima (polygireulima) sp. 4 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 69 subgenus eulima (margineulima) cossmann 1888 eulima (margineulima)? sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 70 genus melanella bowdich 1822 melanella sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 70 family aclididae g.o. sars 1878 genus (graphis) jeffreys 1867 graphis danica ravn 1933 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 70 graphis sp. 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 71 graphis? sp. 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 71 graphis? sp. 3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 72 superfamily triphoroidea gray 1847a family triphoridae gray 1847a subfamily triphorinae gray 1847a genus epetrium harris & burrows 1891 epetrium? cretacea (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 72 epetrium? crassigranulata (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . 72 epetrium? separabilis (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . . . . 73 epetrium? sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 73 genus triphora blainville 1828 subgenus triphora (ogivia) harris & burrows 1891 triphora (ogivia) faxensis ravn 1933 . . . . . . . . . . . . . . . . . . . . . . . . . . 74 triphora (ogivia) sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 74 triphoridae gen. et sp. indet. 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 75 triphoridae gen. et sp. indet. 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 75 family cerithiopsidae h. & a. adams 1853 subfamily cerithiopsinae h. & a. adams 1853 1010 genus cerithiopsis forbes & hanley 1851 cerithiopsis unisulcata ravn 1933 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 76 cerithiopsis aff. unisulcata ravn 1933 . . . . . . . . . . . . . . . . . . . . . . . . . . 76 cerithiopsis rosenkrantzi (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . . 76 cerithiopsis bruennichi ravn 1933 . . . . . . . . . . . . . . . . . . . . . . . . . . . . 77 cerithiopsis sp. 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 77 cerithiopsis sp. 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 78 cerithiopsis sp. 3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 78 cerithiopsis sp. 4 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 78 cerithiopsis sp. 5 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 79 cerithiopsis sp. 6 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 79 cerithiopsis sp. 7 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 79 cerithiopsis sp. 8 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 80 cerithiopsis sp. 9 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 80 cerithiopsis sp. 10 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 80 genus zaclys finlay 1926 zaclys? selandica (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 81 zaclys? nuetzeli n. sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 81 genus eocolina chavan 1952 eocolina sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 82 genus retilaskeya marshall 1978 retilaskeya ravni nützel 1998 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 83 retilaskeya sp. 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 83 retilaskeya sp. 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 84 genus krachia baluk 1975 krachia sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 84 genus specula finlay 1926 specula angustisulcata (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . . . . 84 subfamily aliptinae marshall 1978 genus cerithiopsidella bartsch 1911 cerithiopsidella trinodosa (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . 85 cerithiopsidella sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 85 genus vatopsis gründel 1980 subfamily seilinae golikov & starobogatov 1975 genus variseila dockery 1993 variseila eocostata nützel 1998 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 86 variseila fissicosta (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 86 variseila sp. 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 87 variseila sp. 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 87 genus seila a. adams 1861 seila sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 88 subgenus seila (notoseila) finlay 1926 seila (notoseila) sp. 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 88 seila (notoseila) sp. 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 88 genus thereitis le renard 1997 thereitis tricingulata (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . . . . 89 11 family newtoniellidae korobkov 1955 subfamily newtoniellinae korobkov 1955 genus cerithiella verrill 1882 cerithiella faxensis (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 89 cerithiella fenestrata (ravn 1902b) . . . . . . . . . . . . . . . . . . . . . . . . . . . . 89 cerithiella sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 90 genus trituba jousseaume 1884 trituba obliquecostulata (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . . 90 subfamily eumetulinae golikov & starobogatov 1975 genus eumetula thiele 1912 eumetula multituberculata nützel 1998 . . . . . . . . . . . . . . . . . . . . . . . . 91 eumetula jenseni (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 91 eumetula sp. 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 92 eumetula sp. 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 92 eumetula sp. 3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 92 clade neogastropoda cox 1960 superfamily buccinoidea rafinesque 1815 family fasciolariidae gray 1853 subfamily fusininae wrigley 1927 genus fusinus rafinesque 1815 fusinus sp. 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 93 fusinus sp. 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 93 genus dolicholatirus bellardi 1886 dolicholatirus sp. 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 94 dolicholatirus sp. 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 94 genus conradconfusus snyder 2002 conradconfusus parvus (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . . . 95 conradconfusus subglaber (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . 95 conradconfusus sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 96 superfamily muricoidea rafinesque 1815 family muricidae rafinesque 1815 genus pterynotus swainson 1833 subgenus pterynotus (pterochelus) jousseaume 1880 pterynotus (pterochelus) sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 96 family costellariidae macdonald 1860 genus vexillum röding 1798 vexillum? sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 96 family mitridae swainson 1829 subfamily mitrininae swainson 1829 genus mitra lamarck 1798 mitra subglabra (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 97 mitra glabra (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 97 mitra faxensis (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 98 family volutidae rafinesque 1815 1212 subfamily scaphellinae gray 1857 genus scaphella swainson 1832 scaphella faxensis (ravn 1902b) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 98 family volutomitridae gray 1854 genus conomitra conrad 1865 conomitra sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 98 superfamily conoidea fleming 1822 family clavatulidae gray 1853 genus turricula schumacher 1817 turricula faxensis (ravn 1902b) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 99 turricula pusilla (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 99 superfamily cancellarioidea forbes & hanley 1851 family cancellariidae forbes & hanley 1851 subfamily plesiotritoninae beu & maxwell 1987 genus plesiotriton fischer 1884 plesiotriton steni schnetler & petit 2006 . . . . . . . . . . . . . . . . . . . . . . . . 99 subfamily cancellariinae forbes & hanley 1851 genus unitas palmer 1947 unitas anderseni schnetler & petit 2006 . . . . . . . . . . . . . . . . . . . . . . . 100 unitas aliceae schnetler & petit 2006 . . . . . . . . . . . . . . . . . . . . . . . . . 100 unitas sp. 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 101 unitas sp. 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 101 genus admetula cossmann 1889 admetula rosenkrantzi schnetler & petit 2006 . . . . . . . . . . . . . . . . . . . 101 admetula faksensis schnetler & petit 2006 . . . . . . . . . . . . . . . . . . . . . 101 genus semitriton cossmann 1903 semitriton biplicatus (ravn 1902b) . . . . . . . . . . . . . . . . . . . . . . . . . . . 102 genus tatara fleming 1950 tatara danica schnetler & petit 2006 . . . . . . . . . . . . . . . . . . . . . . . . . 102 clade heterobranchia informal group lower heterobranchia superfamily acteonoidea d'orbigny 1843 family acteonidae d'orbigny 1843 subfamily acteoninae d'orbigny 1843 genus acteon montfort 1810 acteon sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 103 genus rictaxis dall 1871 rictaxis? selandica (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 103 acteonoidea gen. et sp. indet. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 104 superfamily rissoelloidea gray 1850 family rissoellidae gray 1850 genus rissoella gray 1847a rissoella? sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 104 superfamily architectonicoidea gray 1850 family architectonicidae gray 1850 13 genus pseudotorinia sacco 1892 pseudotorinia faxense (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . . . 105 genus nipteraxis cossmann 1916 nipteraxis poulseni (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 105 genus pseudomalaxis fischer 1885 pseudomalaxis sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 105 family orbitestellidae iredale 1917 genus orbitestella iredale 1917 orbitestella sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 106 family amphitomariidae bandel 1996 genus neamphitomaria bandel 1988 neamphitomaria sp. 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 106 neamphitomaria sp. 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 106 neamphitomaria sp. 3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 107 superfamily mathildoidea dall 1889 family mathildidae dall 1889 genus gegania jeffreys 1884 gegania rosenkrantzi (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . . . 107 genus mathilda semper 1865 mathilda unicarinata (ravn 1933) . . . . . . . . . . . . . . . . . . . . . . . . . . . 107 mathilda sp. 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 108 mathilda sp. 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 108 genus acrocoelum cossmann 1888 acrocoelum? sp. 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 109 acrocoelum sp. 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 109 genus clathrobaculus cossmann 1912 clathrobaculus? sp. 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 109 clathrobaculus? sp. 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 110 superfamily pyramidelloidea gray 1840 family pyramidellidae gray 1840 genus odostomia fleming 1813 odostomia sp. 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 110 odostomia sp. 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 110 family amathinidae ponder 1987 genus leucotina a. adams 1860 leucotina sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 111 subfamily syrnolinae saurin 1958 genus puposyrnola cossmann 1921 puposyrnola sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 111 clade opisthobranchia informal group opisthobranchia clade cephalaspidea superfamily philinoidea gray 1850 (1815) family cylichnidae h. & a. adams 1854 genus acteocina gray 1847a 1414 acteocina sp. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 112 incertae sedis incertae sedis sp. 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 112 incertae sedis sp. 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 112 references . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 113 15 abstract lauridsen, b.w. & schnetler, k.i. 2014: a catalogue of danian gastropods from the baunekule facies, faxe formation, denmark. geological survey of denmark and greenland bulletin 32, 117 pp. this catalogue of 194 gastropod taxa is based on the collection of danian gastropods from the baunekule facies, faxe formation in eastern denmark. the gastropod fauna is extremely rich and well preserved. most of the gastropods (106 species) are referred to genus level only, 9 morphotypes to even higher taxonomical levels and 79 gastropods are referred to species level. the gastropods are classified following bouchet & rocroi (2005) as 4 different clades: vetigastropoda (represented by 26 species and 10 superfamilies), caenogastropoda (represented by 142 species and 17 superfamilies), heterobranchia (represented by 23 species and 5 superfamilies) and opisthobranchia (represented by 1 species and 1 superfamily). the new species zaclys? nuetzeli n. sp. is introduced. the faxe formation is recognised as a cold-water coral ecosystem with interfingering smaller bryozoan mounds. the baunekule facies is found in the upper part of the coral mound complex of the faxe formation, where it forms isolated lensoidal bodies in the flanks of some of the coral mounds. it is characterised by a high diversity invertebrate fauna that occurs in weakly consolidated coraldominated floatstone to rudstone. the diagenesis of the baunekule facies is of special significance because a high proportion of the originally aragonite-shelled fauna is preserved by recrystallisation to calcite during early burial diagenesis. most of the gastropods are not known from other parts of the faxe fm. the fauna is very important for comparative evolutionary studies of fossil and modern gastropods on cold-water coral mounds. many of the genera have not previously been recorded from danian strata. none of the gastropod species found in the baunekule facies are known for certain to range below the cretaceous–palaeogene boundary. the fauna is comparable to gastropods found on modern cold-water coral mounds in the north atlantic. the gastropod fauna from the baunekule facies is characterised by a very high diversity of rather small millimetre-sized gastropods with a preference for hard substrates; 63.9% of the taxa belong to the browsing carnivore trophic group, feeding mainly on sedentary animals. surprisingly, the fauna contains some common occurrences of typically warm-water species. the fauna consists mostly of cenozoic genera and up to 87% of the species may be endemic to the cold-water coral ecosystem of the faxe formation. the diverse and rather unusual gastropod fauna from the baunekule facies is undoubtedly linked to the evolution of cold-water coral ecosystems. authors’ addresses b.w.l., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: bwl@geus.dk k.i.s., fuglebakken 14, stevnstrup, 8870 langå, denmark 1616 17 introduction fossil gastropods from the upper cretaceous to danian chalk deposits of northern europe are usually rare due to early diagenetic dissolution of the primarily aragonitic shell. a few gastropods may be found, however, as moulds and casts in cemented hardgrounds reflecting early cementation prior to aragonite dissolution. the presence of gastropods is also indicated indirectly by the common gastropod borings in bivalve or brachiopod shells (surlyk 1972; lauridsen & surlyk 2008; lauridsen et al. 2009; sørensen & surlyk 2011; hansen & surlyk 2014). cold-water corals are also comparatively rare in the fossil record (lauridsen and bjerager 2014), due both to the low preservation potential of the aragonitic shell material and to the fact that coral ecosystems in deep water are a geologically young development. the rich and wellpreserved gastropod fauna in the fossil cold-water coral mound complex of the faxe formation is therefore a rarity and highly significant in understanding the evolution and coexistence of gastropods on the cold-water coral mounds. the faunal compositions of modern cold-water coral mounds are not very well studied, but the few studies published indicate a dominance of gastropods among the associated fauna and some of the families represent the top predators on the cnidarians (reed & mikkelsen 1987; freiwald et al. 2002; reed 2002; mortensen & fosså 2006; henry & roberts 2007; taviani et al. 2009). this catalogue represents the first attempt to comprehensively present the rich gastropod fauna from the faxe formation of eastern denmark and the aim is thus to produce a visual record of the danian fauna from faxe based on unique drafted illustrations and modern photographs and hence to make data on the fauna available for future scientific work. thus, although the genus names have been revised, introduction of new taxa has been postponed, undescribed species are generally recorded in open nomenclature. the species zaclys? selandica (ravn 1933) has been revised, however, since the type material of this species contains two different species; the new species zaclys? nuetzeli n. sp. is thus introduced herein. material the material presented here has been collected in the faxe quarry from weakly consolidated coral-dominated floatstone to rudstone, described as the baunekule facies of the faxe formation by lauridsen et al. (2012). it is characterised by a high diversity invertebrate fauna with preservation of both calcitic and originally aragonite-shelled benthic invertebrates. the gastropods that were drawn in the 1960s and are published here for the first time were collected in the early 20th century by fig. i. key figures in the early collection and classification of the faxe gastropod fauna. professor j.p.j. ravn (left) and professor a. rosenkrantz (right). both photographs are used by permission of the natural history museum of denmark. 1818 j.p.j. ravn, k. brünnich nielsen, c. poulsen and a. rosenkrantz (fig. i). collections were made by søren bo andersen and sten lennart jakobsen in 1972 and later donated to the second author. mrs alice rasmussen and her family from faxe collected more material in 1994 and invested significant effort into preparing, identifying, compiling and organising the fossils in a database (fig. ii). mogens stentoft nielsen, odense and ronald janssen (smf, frankfurt a. m.) also made collections in 1994 and a collection made by erik t.s. christiansen has been donated to faxe museum. a small collection was donated to eivind and birgitte palm, sdr. sejerslev, by alice rasmussen. the material forming the basis for the present catalogue is considerably larger than that used in the monograph published by ravn (1933). the authors visited the geological museum in copenhagen in january 2012 and studied the material from the baunekule facies housed there. the ravn (1933) material was located together with material collected later by g.v. olsen, but the type material for the rosenkrantz drawings was not found. subsequently this material was located in the geological museum by s.l. jakobsen. during the visit, a hitherto unknown series of specimens accompanied by small sketches was found. these were obviously the basis for rosencrantz’s series of drawings, since according to labels on the boxes specimens had been taken out for drawing; several unknown species were identified in this collection. previous work on gastropods from the baunekule facies ravn (1902a, b) monographed the molluscs from the cretaceous of denmark, including the danian and the sediments from the faxe quarry. the bulk of the material from this locality consists of internal moulds and external imprints, but ravn also mentioned specimens of leptomaria niloticiformis and cypraeids with the shells apparently preserved; clearly, the baunekule facies was exposed in the quarry at that time. nielsen (1919) mentioned a gastropod fauna with many small species and compared this fauna with the younger selandian fauna from copenhagen. he was the first to isolate the small specimens by washing and sieving the unconsolidated chalk. ravn (1933) monographed the molluscs from faxe and included material from different facies; he described many new species from the baunekule facies. rosenkrantz collected from these exposures in the early 1930s but the material was not included in the work of ravn (1933). in the 1960s, rosenkrantz obviously intended a publication of this material and a revision of the ravn (1933) monograph, and, most probably also of the selandian of copenhagen and the paleocene of nuussuaq, west greenland. he supervised numerous drawings of molluscs from these localities, made by the illustrators erna nordmann, gunni jørgensen and betty engholm and these drawings along with the serially numbered collections are still housed in the rosenkrantz files at the geological museum of copenhagen. in the folders rosenkrantz arranged the faxe molluscs according to the classification of the time and he used both the fig. ii. alice rasmussen, an amateur geologist at geomuseum faxe, with one of her many discoveries – a block full of nautilids of the species hercoglossa danica. 19 drawings and new photos by christian poulsen. he revised some generic names and gave taxonomical remarks on many drawings, in most cases based on wenz (1938– 1944). some new taxa were also suggested, both at genus and species level, but as they were never published they are all nomina nuda. a few of the drawings had sparse details of localities indicating that they were all collected from the same exposures of the baunekule facies (the locality previously termed ‘ravns næse’; for more details on the locality see lauridsen et al. 2012). rosenkrantz also worked on a series of molluscs, according to the drawings, from the paleocene of nuussuaq and in 1970 he published a few of these drawings and introduced a few new species from greenland. the greater part of the drawings and taxonomic notes of the species from faxe, copenhagen and west greenland, however, were never published by him before his death in 1974. most of the drawings of greenlandic species were subsequently published by kollmann & peel (1983) and schnetler & petit (2010) and some drawings of the species from faxe were published by schnetler et al. (2001), schnetler & petit (2006) and schnetler (2013). new species from the baunekule facies have been established by german authors. bandel & kowalke (1997) published a new species of the family pickworthiidae, bandel (1998) erected a new species of the family anatomidae and nützel (1998) established four new species of the family cerithiopsidae. these authors based their work on small collections of faxe gastropods made by søren bo andersen and donated by claus hedegaard. schnetler et al. (2001) presented a revision of the scissurellidae and anatomidae and established three new species. the family cancellariidae was revised by schnetler & petit (2006) and five new species were established. schnetler & lozouet (2012) erected the new genus and species faxetrochus problematicus, and finally schnetler (2013) introduced the new name eoatlanta ravni for a common gastropod species. geological setting the danian deposits of stevns klint and the faxe quarry represent the latest stages in cool-water carbonate deposition in northern europe which started in the late early cretaceous and lasted nearly 40 million years (fig. iii; thomsen 1995; surlyk 1997). northern europe was situated at palaeolatitudes of 35°n to 50°n during the period (smith et al. 1994) and tropical forms such as reef corals and rudists are generally absent. the nutritional content in the ocean was ideal for the development of enormous blooms of coccolithophoroid algae. in the danish area during the late cretaceous the deposition of coccolithic ooze took place in the deeper part of the danish basin and a specially adapted millimetre-size benthic fauna is associated with this environment (surlyk 1972). a highly diverse boreal fauna of bryozoans, bivalves, brachiopods, echinoderms and serpulids thrived in the shallower parts along the margins of the basin (lauridsen & surlyk 2008). fossil gastropods are generally rare in the chalk deposits, probably due to low preservation potential. in the early danian, large asymmetric bryozoan mounds were formed and migrated towards the southeast, parallel to the coastline and towards a nutrientrich current (thomsen 1976; surlyk 1997; surlyk et al. 2006; bjerager & surlyk 2007a, b; nielsen et al. 2009). the mounded bryozoan limestones are referred to the stevns klint formation and span from the middle part of the early danian to the boundary between the midand late danian (surlyk et al. 2006). from the mid-danian, a low diversity but abundant fauna of azooxanthellate corals constructed biogenic mounds along the easternmost rim of the ringkøbing-fyn high in the faxe area (floris 1980; bernecker & weidlich 1990, 2005; willumsen 1995; lauridsen et al. 2012; lauridsen & bjerager 2014). the faxe formation is defined as a distinct mappable lithostratigraphic unit of interfingering coral and bryozoan limestone passing laterally into bryozoan limestones of the stevns klint formation (fig. iv) (lauridsen et al. 2012). the mounds are formed predominantly by the frame-building corals dendrophyllia candelabrum, with minor occurrences of oculina becki and faksephyllia faxensis. the mounds are the result of complex interactions between biological and geological processes (lauridsen & bjerager 2014). the mound 2020 ecosystem consists of numerous individual mounds of 50–100 m in diameter. smaller bryozoan mounds and intervals with an octocoral-rich facies interfinger with the larger coral mounds (lauridsen et al. 2012). the low diversity of frame-building corals and the lack of algal borings and shallow-water sedimentary structures suggest that the corals grew in relatively deep water below the photic zone. the presence of a diverse and abundant stylasterine fauna on the fossil mounds suggests a stable palaeoenvironment, probably in a bathymetric depth range of 200–400 m (lauridsen & bjerager 2014). the danian cold-water coral mounds show strong similarities to the coral mound ecosystems encountered at high latitudes in the deep waters of the north atlantic today. the term ‘cold-water coral mounds’ used for the modern ecosystem is therefore also applied to the coral mounds of the faxe formation. notes on the locality and the baunekule facies a wide range of coral limestone facies has been recorded in the faxe formation. the dominant biogenic moundbuilding facies with essentially in-place fossils mainly comprises coral rudstone to floatstone and bafflestone (lauridsen et al. 2012). associated facies include a wide range of fine to coarse grainstones, packstones and wackestones. a distinct facies has been recorded from the upper part of the faxe formation, composed of a weakly consolidated coral limestone with an unusually well-preserved, high-diversity invertebrate fauna; this is termed the baunekule facies (see below; lauridsen et al. 2012). the intercalated bryozoan limestone facies consists of rudstone, floatstone, packstone, and wackestone. the degree of diagenesis varies throughout the formation from extensively diagenetically altered to almost pristine coral limestone. early diagenesis was characterised by dissolution of aragonite skeletons and associated calcite deposition and precipitation of matrix cement, cement ? land in the danian sea in the danian fig. xx3 fig. iii. a palaeogeographic reconstruction of the danian sea in north europe. the position of the faxe quarry is indicated by a star. modified from lauridsen & damholt (2011). 21 rims, and interparticle and intraparticle replacement cements (bernecker & weidlich 1990; willumsen 1995; bjerager et al. 2010). later diagenesis involved recrystallisation of shells and cementation of the limestones. the baunekule facies has been encountered at various times during the progressive expansion of the quarry; two important localities with respect to the collections described here were the baunekule locality (exposed from around 1900 to the 1930s) and the stationsvej locality (exposed in the 1970s and 1990s; for more information on the localities, see lauridsen et al. 2012 and fig. iv). the rosenkrantz collection on which the drawings are based was collected at the baunekule locality; all other collections are from the stationsvej locality (fig. iv). the baunekule facies occurs in the upper part of the coral mound complex of the faxe formation, where it forms isolated lensoidal bodies within the flanks of some of the coral mounds (lauridsen et al. 2012). fig. iv. photograph and accompanying sketch of coral mounds intercalating with bryozoan mounds in the faxe quarry at the type locality of the faxe formation, close to geomuseum faxe (gmf); two persons for scale are encircled in front of the museum. the flint layers in the bryozoan limestone emphasise the original bedding of the mounds. octocoral limestones are rare in the quarry and are commonly found to have accumulated prior to scleractinian coral growth. the coral limestone at the type locality is predominantly of the species dendrophyllia candelabrum but minor patches of faksephyllia faxensis are also recorded. the two localities (stationsvej and baunekule) where the baunekule facies has been recorded previously are indicated on the accompanying map. modified from lauridsen & bjerager (2014). bryozoan limestone with numerous flint bands s n coral limestone with scleractinian corals octocoral limestone interfingering boundary sweden sweden denmark 100 km germany germany copenhagen stevns faxe gmf s n 500 m 25 m 1 n stationsvej loc. baunekule loc. type loc. faxe fm faxe quarry 8˚ 10˚ 12˚e 57˚n 56˚ 55˚ 2222 the frame-building corals have a bushy growth form and were mound forming. they provided excellent habitats and feeding grounds for a large and diverse benthic invertebrate fauna. brachiopods, bivalves, echinoids, asteroids, serpulids, solitary corals, octocorals and moulds of decapods and gastropods are recorded everywhere in the faxe formation (fig. v). however, the degree of preservation is extremely variable. more than 25  000 invertebrate fossils in the alice rasmussen collection at geomuseum faxe were collected from the baunekule facies at the stationsvej locality, representing at least 300 species. the fossils represent a wide range of modes of life and trophic levels being represented by brachiopods, bivalves, serpulids, echinoderms and corals (lauridsen et al. 2012). the corals are dominated by species of the octocoral moltkia and by scleractinian solitary coral species of parasmilia. originally aragonitic stylasterine hydrocorals are also rather common (lauridsen & bjerager 2014). the high density and diversity of the originally aragonite-shelled fauna provide a rare taphonomic window into the past ecosystem. this is exemplified by the high abundance of millimetre-sized gastropods representing 194 different species. in other parts of the faxe formation, the small gastropods are only found as rare moulds or casts and are often very difficult to identify to species or even to genus level (fig. vi). fig. v. the rich associated fauna of the faxe formation is dominated by originally calcitic, shelly invertebrates. a: moulds of the frame-building coral dendrophyllia candelabrum. b: two specimens of the bivalve chlamys hennigi are well hidden between the coral branches. c: bryozoans are also common. d: mould of the decapod dromiopsis rugosa which is common throughout the cold-water coral limestone. e: echinoid plates from a cidaris sp. small specimens are common, being specially adapted to live in between the coral branches. f: mould of a solitary coral. scale bar is 10 mm. 20 mm fig. vi. a large specimen of the gastropod genus leptomaria encountered in the coral limestone of the faxe formation. photograph: jesper milàn, geomuseum faxe. 23 remarks on the fauna preservation the diagenesis of the baunekule facies is of special importance as early burial diagenesis resulted in preservation of many of the aragonitic faunal elements by recrystallisation to calcite prior to the precipitation of high mg-calcite cement (lauridsen et al. 2012). the carbonate mud matrix is only weakly cemented and the macro and microfossils are accordingly easy to prepare out, in contrast to the fossils from the lithified part of the faxe formation. many of the more common gastropods are preserved both as gastropod shells and as moulds and imprints. the latter preservation types are comparable to the gastropod material encountered in other parts of the faxe formation. bøggild (1930) studied the structures of mollusc shells and found that the specimens were preserved with the exterior shell and that the primary aragonite was recrystallised into grainy calcite. this preservation allows study of the sculpture and the protoconch. the bulk of the material consists of small species and juvenile specimens of larger species; species larger than 10 mm are almost completely absent. the collection of specimens from the baunekule facies has been kept in small labelled glass tubes since the collection of material in the early 1900s. in some cases, however, shells have deteriorated probably as a result of chemical changes linked to the glass tubes or to varying humidity; calcite crystals on the shells or dissolution features can be observed on specimens in the collection. number of taxa j.p.j. ravn was the first to work systematically with the gastropods from the faxe formation. he recorded 56 mollusc species of which 27 were gastropods (ravn 1902b). later, ravn (1933) identified 137 mollusc species, of which 98 species were gastropods. rosenkrantz classified the gastropods and bivalves from faxe into taxa labelled 9–267; although rosenkrantz later revised his early classification, the data were never published. the present study has resulted in the identification of 194 species of gastropods (table 1); most of the species are in open nomenclature and some cannot be determined even to genus level. stratigraphic ranges of the genera the stratigraphic ranges of the genera encountered in the baunekule facies are listed in table 2. the distributions are mainly based on wenz (1938–44), rosenkrantz (1960), knight et al. (1960), pacaud et al. (2000) and kiel (2001). the present study lists 194 species of gastropods, representing 104 different genera. many of the genera have not been recorded from danian strata prior to this study: one genus was previously only known from the jurassic, one from the jurassic to cretaceous, two from the cretaceous, and three exclusively from danian strata. forty genera are known from the cretaceous to the cenozoic, spanning the cretaceous–palaeogene (k/ pg) boundary, 49 genera were only known from younger cenozoic strata, and six genera were previously only known from the recent. none of the gastropod species found in the baunekule facies is known to range below the k/pg boundary. comparison with other paleocene gastropod faunas comparisons of gastropod species from faxe with other lower paleocene strata are complicated. however, in table 3 some of the relevant european successions have been compiled and listed. these are the lower danian cerithium limestone member (rødvig formation) from denmark, the middle to upper danian of vigny, france, the danian deposits of nuussuaq, west greenland and the selandian lellinge greensand from the copenhagen area, denmark. it is apparent from these comparisons that only two of the localities have species in common with the baunekule facies. these are the cerithium limestone member, which predates the baunekule facies in the danish basin and the danian outcrops of vigny which in part represent a similar palaeoenvironment (a cold-water coral ecosystem). 2424 1 fissurellidae rimula? sp. rc hr/cb 2 fissurellidae emarginula coralliora lundgren 1867 vc hr/cb 3 leptomaria niloticiformis (von schlotheim 1820) vc cb 4 leptomaria sp. vr cb 5 temnotropidae temnotropis sp. r hr 6 porcellidae faxetrochus problematicus schnetler & lozouet 2012 vr cb? 7 scissurellidae scissurella annulata ravn 1933 vc cb 8 scissurellidae scissurella (s. l.) aliceae schnetler, lozouet & pacaud 2001 c cb 9 scissurellidae scissurella ravni (schnetler, lozouet & pacaud 2001) r cb 10 anatomidae anatoma hedegaardi (bandel 1998) r cb 11 anatomidae anatoma rosenkrantzi schnetler, lozouet & pacaud 2001 vr cb 12 trochidae osilinus carinatus (ravn 1933) a cb 13 trochidae jujubinus sp. vr cb 14 trochidae stomatella sp. vr cb 15 turbinidae leucorhynchia marginata ravn 1933 vr cb 16 turbinidae skeneoides sp. r cb 17 ataphridae ataphrus sp. vr cb 18 margaritidae margarites bruennichi (ravn 1933) a cb 19 angariidae angaria depressa (ravn 1933) vc cb 20 colloniidae collonia (circulopsis) pusilla (ravn 1933) vc cb 21 colloniidae vexinia sp. vr cb 22 eucycloscalidae eucycloscala ultima (ravn 1933) rc cb 23 chilodontidae danilia faxensis (ravn 1933) a hm 24 chilodontidae danilia quadricordata (ravn 1933) a hm 25 chilodontidae danilia fenestrata (ravn 1933) a hm 26 chilodontidae chilodontinae gen. et sp. indet. vr hm 27 cerithiidae ataxocerithium faxensis (ravn 1933) vc hp 28 cerithiidae ataxocerithium sp. 1 vc hp 29 cerithiidae ataxocerithium sp. 2 vc hp fig. no. species frequency† palaeoecology§ family table 1. taxa encountered in the study, listed with frequency and inferred feeding strategies 30 cerithiidae bittium transenna (bayan 1873)? rc hp 31 metacerithiidae metacerithium? sp. vr hp 32 campanilidae campanile? pseudotelescopium (ravn 1902) rc hp 33 campanilidae campanile? subglabra (ravn 1933) r hp 34 campanilidae campanile? sp. 1 vr hp 35 campanilidae campanile? sp. 2 vr hp 36 trypanaxidae trypanaxis faxensis ravn 1933 rc hp 37 turritellidae mesalia? sp. 1 r su 38 turritellidae mesalia? sp. 2 r su 39 turritellidae mesalia? sp. 3 vr su 40 siliquariidae tenagodus ornatus (lundgren 1867) c su 41 siliquariidae tenagodus sp. vr su pleurotomariidae pleurotomariidae † a: abundant (> 500 specimens). vc: very common (100–499 specimens). c: common (50–99 specimens). rc: rather common (20–49 specimens). r: rare (5–19 specimens). vr: very rare (1–4 specimens). § the palaeoecological classification is adopted from todd (2000). cb: browsing carnivores. cp: predatory carnivores. hm: herbivores on fine-grained substrates. ho: herbivorous omnivores. hp: herbivores on algal substrates. hr: herbivores on rock, rubble or coral substrates. su: suspension feeders. 42 capulidae trichotropis? sp. vr su 25 fig. no. species frequency† palaeoecology§ family table 1. (continued) 44 cypraeidae palaeocypraea spirata (von schlotheim 1820) c ho 45 cypraeidae palaeocypraea sp. vr ho 46 cypraeidae bernaya (protocypraea) globuliformis (ravn 1902) vr ho 47 eocypraeidae eocypraea danica (schilder 1928) rc ho 49 pickworthiidae sansonia hedegaardi bandel & kowalke 1997 c hr/su ? 50 pickworthiidae sansonia sp. r hr/su ? 51 pickworthiidae mareleptopoma? sp. r hr/su ? 52 pickworthiidae urceolabrum sp. 1 vr hr/su ? 53 pickworthiidae urceolabrum sp. 2 vr hr/su ? 54 pickworthiidae faxia macrostoma ravn 1933 c hr/su ? 55 naticidae naticidae gen. et sp. indet. vr cp 56 rissoidae zebina sp. 1 r hp 57 rissoidae zebina sp. 2 vr hp 59 rissoidae rissoinidae gen. et sp. indet. rc hp 60 caecidae caecum sp. vr hr/hp 61 tornidae circulus sp. 1 vr hr/hp 62 tornidae circulus sp. 2 vr hr/hp 63 tornidae teinostoma glaberrimum ravn 1933 r hr/hp 64 tornidae sigaretornus sp. vr hr/hp 65 ranellidae ranella faxensis ravn 1933 r cp 66 ranellidae sassia faxense (ravn 1933) c cp 67 ranellidae sassia sp. vr cp 68 ranellidae cymatium (monoplex) subglabrum (ravn 1902) r cp 69 elachisinidae laeviphitus sp. vr ? 70 hipponicidae hipponixsp. r hr/su 71 hipponicidae eoatlanta ravni schnetler 2013 vc hr/su ? 72 triviidae johnstrupia faxensis ravn 1933 vr cb 73 epitoniidae cerithiscala tricincta (ravn 1933) rc cb 74 epitoniidae cerithiscala sp. 1 r cb 75 epitoniidae cerithiscala sp. 2 vr cb 76 epitoniidae opalia sp. rc cb 77 epitoniidae acrilla elegans (ravn 1902) r cb 78 epitoniidae acrilla sp. 1 vr cb 79 epitoniidae acrilla sp. 2 vr cb 80 epitoniidae acrilla sp. 3 vr cb 81 epitoniidae acrilla? sp. vr cb 82 epitoniidae cirsotrema sp. vr cb 83 eulimidae eulima (polygireulima) danica ravn 1933 rc cb 84 eulimidae eulima (polygireulima) sp. 1 vr cb 85 eulimidae eulima (polygyreulima) sp. 2 vr cb 86 eulimidae eulima (polygyireulima) sp. 3 vr cb 87 eulimidae eulima (polygireulima) sp. 4 vr cb 88 eulimidae eulima (margineulima?) sp. vr cb 89 eulimidae melanella sp. vr cb 90 aclididae graphis danica (ravn 1933) rc cb 48 littorinidae littoraria (littorinopsis) faxensis (ravn 1933) hrr 43 capulidae capulidae gen. et sp. indet. vr su? 58 rissoidae pseudotaphrus sp. vr hp? 2626 fig. no. species frequency† palaeoecology§ family table 1. (continued) 91 aclididae graphis sp. 1 r cb 92 aclididae graphis? sp. 2 vr cb 93 aclididae graphis? sp. 3 vr cb 94 triphoridae epetrium? cretacea (ravn 1933) a cb 98 triphoridae triphora (ogivia) faxensis ravn 1933 vr cb 99 triphoridae triphora (ogivia) sp. r cb 100 triphoridae triphoridae gen. et sp. indet. 1 vr cb 101 triphoridae triphoridae gen. et sp. indet. 2 r cb 102 cerithiopsidae cerithiopsis unisulcata ravn 1933 vc cb 103 cerithiopsidae cerithiopsis aff. unisulcata ravn 1933 vr cb 104 cerithiopsidae cerithiopsis rosenkrantzi (ravn 1933) r cb 105 cerithiopsidae cerithiopsis bruennichi ravn 1933 rc cb 106 cerithiopsidae cerithiopsis sp. 1 vr cb 107 cerithiopsidae cerithiopsis sp. 2 vr cb 108 cerithiopsidae cerithiopsis sp. 3 vr cb 109 cerithiopsidae cerithiopsis sp. 4 r cb 110 cerithiopsidae cerithiopsis sp. 5 vr cb 111 cerithiopsidae cerithiopsis sp. 6 vr cb 112 cerithiopsidae cerithiopsis sp. 7 vr cb 113 cerithiopsidae cerithiopsis sp. 8 r cb 114 cerithiopsidae cerithiopsis sp. 9 vr cb 115 cerithiopsidae cerithiopsis sp. 10 vr cb 116 cerithiopsidae zaclys? selandica (ravn 1933) rc cb 117 cerithiopsidae zaclys? nuetzeli n. sp. r cb 118 cerithiopsidae eocolina sp. r cb 119 cerithiopsidae retilaskeya ravni nützel 1998 rc cb 120 cerithiopsidae retilaskeya sp. 1 r cb 121 cerithiopsidae retilaskeya sp. 2 vr cb 122 cerithiopsidae krachia sp. r cb 123 cerithiopsidae specula angustisulcata (ravn 1933) vc cb 124 cerithiopsidae cerithiopsidella trinodosa (ravn 1933) r cb 125 cerithiopsidae cerithiopsidella sp. vr cb 126 cerithiopsidae vatopsis metaxiformis nützel 1998 c cb 127 cerithiopsidae variseila eocostata nützel 1998 r cb 128 cerithiopsidae variseila fissicosta (ravn 1933) c cb 129 cerithiopsidae variseila sp. 1 vr cb 130 cerithiopsidae variseila sp. 2 vr cb 131 cerithiopsidae seila sp. vr cb 132 cerithiopsidae seila (notoseila) sp. 1 rc cb 133 cerithiopsidae seila (notoseila) sp. 2 vr cb 134 cerithiopsidae thereitis tricingulata (ravn 1933) rc cb 135 newtoniellidae cerithiella faxensis (ravn 1933) c cb 136 newtoniellidae cerithiella fenestrata (ravn 1902) r cb 137 newtoniellidae cerithiella sp. r cb 138 newtoniellidae trituba obliquecostulata (ravn 1933) c cb 95 triphoridae epetrium? crassigranulata (ravn 1933) rc cb 96 triphoridae epetrium? separabilis (ravn 1933) r cb 97 triphoridae epetrium? sp. vr cb 27 fig. no. species frequency† palaeoecology§ family table 1. (continued) 139 newtoniellidae eumetula multituberculata nützel 1998 rc cb 140 newtoniellidae eumetula jenseni (ravn 1933) c cb 146 fasciolariidae dolicholatirus sp. 1 r cp 147 fasciolariidae dolicholatirus sp. 2 vr cp 148 fasciolariidae conradconfusus parvus (ravn 1933) vc cp 149 fasciolariidae conradconfusus subglaber (ravn 1933) vc cp 150 fasciolariidae conradconfusus sp. rc cp 151 muricidae pterynotus (pterochelus) sp. r cp 152 costellariidae vexillum sp. r cp 153 mitridae mitra subglabra (ravn 1933) vc cp 154 mitridae mitra glabra (ravn 1933) r cp 155 mitridae mitra faxensis (ravn 1933) r cp 156 volutidae scaphella faxensis (ravn 1902) vr cp 157 volutomitridae conomitra sp. r cp 158 clavatulidae turricula faxensis (ravn 1902) r cp 159 clavatulidae turricula pusilla (ravn 1933) r cp 160 cancellariidae plesiotriton steni schnetler & petit 2006 r cb 161 cancellariidae unitas anderseni schnetler & petit 2006 r cb 162 cancellariidae unitas aliceae schnetler & petit 2006 r cb 163 cancellariidae unitas sp. 1 vr cb 164 cancellariidae unitas sp. 2 r cb 165 cancellariidae admetula rosenkrantzi schnetler & petit 2006 vr cb 166 cancellariidae admetula faksensis schnetler & petit 2006 r cb 167 cancellariidae semitriton biplicatus (ravn 1902) rc cb 168 cancellariidae tatara danica schnetler & petit 2006 rc cb 169 acteonidae acteon sp. r cp 170 acteonidae rictaxis? selandica (ravn 1933) c cp 171 acteonidae? acteonoidea, gen. et sp. indet. rc cb 172 rissoellidae rissoella sp. vr cb? 173 architectonicidae pseudotorinia faxense (ravn 1933) vc cb 174 architectonicidae nipteraxis poulseni (ravn 1933) r cb 175 architectonicidae pseudomalaxis sp. c cb 176 architectonicidae orbitestella sp. vr cb 177 amphitomariidae neamphitomaria sp. 1 vr cb 178 amphitomariidae neamphitomaria sp. 2 vr cb 179 amphitomariidae neamphitomaria sp. 3 vr cb 180 mathildidae gegania rosenkrantzi (ravn 1933) c cb 181 mathildidae mathilda unicarinata (ravn 1933) rc cb 182 mathildidae mathilda sp. 1 vr cb 183 mathildidae mathilda sp. 2 vr cb 184 mathildidae acrocoelum? sp. 1 vr cb 185 mathildidae acrocoelum? sp. 2 vr cb 186 mathildidae clathrobaculus? sp. 1 vr cb 142 newtoniellidae eumetula sp. 2 vr cb 143 newtoniellidae eumetula? sp. 3 vr cb 144 fasciolariidae fusinus sp. 1 r cp 145 fasciolariidae fusinus sp. 2 vr cp 141 newtoniellidae eumetula sp. 1 r cb 2828 fig. no. species frequency† palaeoecology§ family table 1. (continued) 187 mathildidae clathrobaculus? sp. 2 r cb 188 pyramidellidae odostomia sp. 1 vr cb 190 amathinidae leucotina sp. r cb 191 amathinidae puposyrnola sp. vr cb 192 cylichnidae acteocina sp. r cp 193 incertae sedis sp. 1 vr ? 194 incertae sedis sp. 2 vr ? 189 pyramidellidae odostomia sp. 2 c cb palaeoecology the coral branches on the cold-water coral mounds served as substrate, shelter and food for a rich and diverse associated fauna, many of the species being especially adapted for a life on the coral mounds. the rich and unusual gastropod fauna presented here is ideal for palaeoecological studies. however, an in-depth study is beyond the scope of this paper, so the subject is only discussed briefly below. the classification of gastropod feeding characteristics used in this study is taken from todd (2000; table 1). the functional morphology of gastropod shells is complicated, mainly because the shells of the gastropods are less associated with the substrate compared to bivalves. in table 1 the frequencies of the species in terms of number of individuals and their presumed diets are listed, and these data are plotted in figs vii and viii. the six most abundant species represented by more than 500 specimens are: margarites bruennichi (ravn 1933), osilinus carinatus (ravn 1933), epetrium? cretacea (ravn 1933), danilia faxensis (ravn 1933), danilia quadricordata (ravn 1933) and danilia fenestrata (ravn 1933). the dominant trophic group in terms of number of species (fig. viii) is the browsing carnivores (124 taxa, 63.9% of the gastropod fauna), feeding on sedentary animals such as sponges and corals, followed by the predatory carnivores feeding on mobile macro organisms (21 species, 10.8% of all species). fourteen species (7.2% of all species) are herbivores whereas 3–4% of all species are herbivores on microalgae or suspension feeders, some living stationary in sponges such as tenagodus ornatus of the siliquariidae. three species or genera belong to unknown trophic groups. it is evident from the preliminary results that the gastropod fauna is dominated by species that had a preference for hard substrates, such as the coral mounds. this is unusual and makes the fauna from faxe unique among the described paleocene gastropod faunas which are normally dominated by soft substrate faunas. the faxe fauna represents a thanatocoenosis and this could explain the rather common occurrence of herbivores, possibly representing a fauna from shallower water depths. furthermore, some of the common species, such as members of cypraidae are traditionally regarded as being typically warm-water species; this is anomalous in this deep and inferred cold-coral ecosystem. in summary, the fossil gastropod fauna from baunekule facies is very diverse but is dominated by forms with a preference for hard substrates; 63.9 % of the species belong to the browsing carnivores trophic group. surprisingly, the fauna contains some common occurrences of typically warm-water species. the fauna consists mainly of cenozoic genera and up to 87% of the species may be endemic to the cold-water coral ecosystem of the faxe formation; substantiation of this point demands more work on the contemporaneous cemented hardgrounds occurring in the bryozoan limestone of the rødvig formation. 29 table 2. known stratigraphic ranges of genera recorded from the baunekule facies, faxe formation jurassic clathrobaculus jurassic to cretaceous temnotropis jurassic to recent trichotropis cretaceous ataphrus neamphitomaria cretaceous to cenozoic acrilla acteocina acteon angaria ataxocerithium bernaya (protocypraea) campanile cerithiella cerithiopsis collonia conradconfusus danilia emarginula eocypraea eumetula eucycloscala fusinus gegania hipponix jujubinus leptomaria littoraria (littorinopsis) mathilda mesalia notoseila palaeocypraea pseudomalaxis pseudotorinia rimula seila surcula teinostoma tenagodus thereitis unitas urceolabrum variseila vatopsis cenozoic acrocoelum admetula anatoma bittium caecum cerithiopsidella cerithiscala circulus cirsotrema conomitra cymatium (monoplex) dolicholatirus eoatlanta eocolina epetrium eulima (polygireulima) eulima (margineulima) graphis krachia leucorhynchia mareleptopoma margarites mitra nipteraxis odostomia omalaxis osilinus plesiotriton pliciscala pseudotaphrus pterynotus (pterochelus) puposyrnola ranella retilaskeya rictaxis sansonia scissurella seila (notoseila) semitriton specula stormatella tatara trituba triphora (ogivia) trypanaxis vexillum vexinia zaclys zebina danian exclusively faxetrochus faxia johnstrupia recent laeviphitus leucotima orbitestella rissoella skeneoides sigaretornus 3030 genera species faxe, denmark (baunekule facies, faxe formation) middle danian cold-water coral mounds 194 – – this bulletin stevns, denmark (cerithium limestone member, rødvig formation) early danian open marine carbonate shelf 62 admetula anatoma cerithiella cerithiopsis cerithiscala dolicholatirus rissoina tatara thereitis unitas vatopsis vexillum eulima danica campanile pseudotelescopium leptomaria niloticiformis/ leptomaria meyeri sassia faxense rosenkrantz (1940) hansen (unpublished paper) nuussuaq, west greenland (agatdal formation) danian cold-water coral thickets 252 acrocoelum admetula cerithiella cerithiopsis circulus conomitra eocypraea fusinus gegania leptomaria mathilda odostomia palaeocypraea scissurella teinostoma turricula unitas none rosenkrantz (1970) kollmann & peel (1983) pacaud & schnetler (1999) schnetler & petit (2010) vigny, france (vigny limestones) middle–late danian open marine, warm shallowwater(?) possible cold-water corals are present 83 emarginula coralliora osilinus carinatus* eucycloscala ultima** cerithiopsis bruennichi mitra glabra mitra subglabra pacaud et al. (2000) montenat et al. (2002) copenhagen, denmark (lellinge greensand) selandian deep shelf/inlet, low-energy 125 none none koenen (1885) ravn (1939) schnetler (2001) * previously boutillieria carinata ** previously urceolabrum ultimum locality/stratigraphy referencesin common with baunekule faciesno. of species palaeoenvironment age table 3. comparison of baunekule species and genera with other paleocene gastropod faunas 31 gastropods on modern cold-water coral mounds this study suggests that most of the gastropods were endemic to the danian cold-water coral ecosystem and it is therefore relevant to compare the occurrences with data from modern cold-water coral ecosystems. data on gastropods from modern cold-water ecosystems, however, are rather limited and only very few studies have worked with the gastropods in any detail (reed & mikkelsen 1987 (oculina mounds in offshore florida); freiwald et al. 2002 (sula reef norwegian shelf); reed 2002 (oculina reef in offshore florida); mortensen & fosså 2006 (mid-norwegian shelf); henry & roberts 2007 (porcupine seabight, atlantic ocean); taviani et al. 2009 (mediterranean)). reed & mikkelsen (1987) studied the molluscs from the modern actively growing oculina mounds offshore florida. the coral species oculina is also represented in faxe but is one of the subordinate frame-building corals. the study lists a total of 230 species-level taxa consisting of 155 gastropods, 68 bivalves, 1 scaphopod, 5 polyplacophorans, and 1 cephalopod. the authors concluded that the carnivorous, mainly coral-eating, molluscs dominated markedly at their deepest locality (around 80 m) where they formed 62.1% of the individuals (reed & mikkelsen 1987). this observation is compatible with a vc c rc r vr 6 15 15 23 52 83 0 10 20 30 40 50 60 70 80 90 n um be r o f s pe cie s fig. xx7 fig. vii. histogram of the frequencies of different species encountered in the studied material. the six most abundant species are represented by more than 500 specimens whereas 83 very rare species are represented by only 1–4 specimens. a: abundant (>500 specimens). vc: very common (100–499 specimens). c: common (50–99 specimens). rc: rather common (20–49 specimens). r: rare (5–19 specimens). vr: very rare (1–4 specimens). fig. viii. feeding strategies of the gastropods. note that the majority of species are carnivorous browsers possibly due to the dominance of sedentary prey. cb: browsing carnivores. cp: predatory carnivores. hm: herbivores on fine-grained substrates. ho: herbivorous omnivores. hp: herbivores on algal substrates. hr: herbivores on rock, rubble or coral substrates. su: suspension feeders. cp hp hr/su?su hohr/hp hm hr hr/cb hr/su ? 21 cb 124 14 67 46 3 2 2 1 2 0 20 40 60 80 100 120 140 n um be r o f s pe cie s 64% 11% 8% 3% 3% 2%4% 1.5% 1% 1% 0.5% 1% fig. xx8 3232 the dominance of browsing carnivores in the baunekule facies. in the study of the recent lophelia mounds in norway, a total of 361 invertebrate species were recorded of which only 25 are gastropods (mortensen & fosså 2006). henry & roberts (2007) studied the lophelia and madrepora mounds of the porcupine seabight in the atlantic ocean and identified 349 invertebrate species of which 47 species are molluscs. the gastropod predators of modern lophelia-madrepora mounds in the mediterranean were studied by taviani et al. (2009). three species from the family muricidae were identified and one of them was observed attached to the corals, whereas two other species only showed indirect relationships. the muricidae, pterynotus (pterochelus) sp., which is represented by a few specimens in the baunekule facies, probably had a similar feeding strategy to its modern relatives. at the present day, the gastropod genus pedicularia is commonly symbiotic on stylasterine corals (zibrowius & cairns 1992; braga-henriques et al. 2011). the association between pediculariids and stylasterine corals has been reported in 12 host species (braga-henriques et al. 2011). specimens of pedicularia are well adjusted to the morphology of their host and show a high degree of plasticity (braga-henriques et al. 2011). all large specimens of stylasterine corals from the baunekule facies have been checked for traces of symbiotic gastropods (lauridsen & bjerager 2014). however, no convincing traces were found, though this could be due to poor preservation potential of the traces. in addition, a rich epifauna of serpulids most likely settled after the death of the corals, covering any symbiotic traces. no fossil pedicularia has been found in the baunekule facies. concluding remarks and future studies the highly diverse gastropod fauna from the baunekule facies is an unusual fauna and undoubtedly linked to the evolution of cold-water coral ecosystems. the gastropods are specially adapted to life on the coral mounds and are dominated by browsing carnivores that preferred hard substrates. the fauna consists mainly of cenozoic genera. more work is needed to study the distribution of gastropods on cold-water coral mounds in time and space. our aim in publishing this first attempt at a modern systematic overview of the gastropod fauna from faxe is to contribute to this long-term goal. acknowledgements this study forms part of a geocenter denmark project led by morten bjerager. financial support from the carlsberg foundation is gratefully acknowledged by bwl. the university of copenhagen supported financially a sixmonth period as guest researcher (bwl) at senckenberg am meer, wilhelmshaven, germany. the following persons are thanked for valuable assistance with locating the fossil material, and with scanning and photographical work: sten lennart jakobsen and morten lunde nielsen (geological museum, university of copenhagen), søren bo andersen and erik thomsen (aarhus university), ane elise schrøder (geomuseum faxe), jørgen kystol (geus) and jan adolfssen (ministry of industry and mineral resources, greenland). thomas hansen (previously of the geological museum) is thanked in particular for help in the museum, taxonomical advice and discussion, and for allowing us access to unpublished data from a submitted manuscript (gastropods from the cretaceous–paleogene boundary in denmark). pierre lozouet (paris), jean-michel pacaud (paris), anders warén (stockholm), andrzej kaim (warsaw) and thomas a. darragh (melbourne, australia), steffen kiel (munich) and alexander nützel (munich) are thanked for help with taxonomic problems, fruitful discussions and advice. alexander nützel furthermore placed photos of the holotype of retilaskeya ravni nützel 33 1998 and cerithiopsis rosenkrantzi at our disposal and helped with photos. jesper milàn (faxe), ronald janssen (frankfurt a. m.), alice rasmussen (faxe, deceased), søren bo andersen (aarhus), sten lennart jakobsen (copenhagen), erik t.s. christiansen (svendborg), eivind and birgitte palm (sdr. sejerslev) and mogens stentoft nielsen (odense) kindly placed their collections at our disposal. the manuscript benefited significantly from the constructive comments of the referees j. alistair crame (british antarctic survey, uk) and thomas hansen (akvaplan-niva, norway). 3434 35 systematic palaeontology abbreviations and classification arf alice rasmussen collection, faxe, denmark (now housed in geomuseum faxe under østsjællands museum, denmark). specimens mentioned have the prefix øsm-10059. ebp eivind and birgitte palm collection, sdr. sejerslev, denmark. ecs erik t.s. christiansen collection, svendborg (now housed in geomuseum faxe, denmark). specimens mentioned have the prefix øsm 10061. gf geomuseum faxe, faxe, denmark gm acronym for registered material in the geological museum, copenhagen, denmark. gpium geologisch-paläontologisches institut und museum, hamburg, germany. isl kai ingemann schnetler collection, langå, denmark. jmp jean-michel pacaud collection, paris, france. mguh geological museum type collection, copenhagen, denmark. mnhn muséum national d’histoire naturelle, paris, france. mno mogens stentoft nielsen collection, odense, denmark. mguh geological museum type collection, natural history museum of denmark, university of copenhagen, denmark. m.u.h. museum universitas hafniensis (museum of the university of copenhagen), unpublished museum names (nomina nuda). øsm østsjællands museum, store heddinge, denmark rgm naturalis biodiversity center (palaeontology department), leiden, the netherlands (formerly rijksmuseum van geologie en mineralogie). rkz acronym for specimens illustrated in the rosenkrantz files of drawings, in the geological museum, copenhagen, denmark. sgpih geologisch-paläontologisch institut und museum der universität hamburg, germany. smf forschungsinstitut senckenberg, frankfurt am main, germany. the gastropods are arranged in accordance with the family-level classification of bouchet & rocroi (2005). four different clades are represented: vetigastropoda, caenogastropoda, heterobranchia and opisthobranchia. the clade caenogastropoda contains the clades sorbeoconcha and hypsogastropoda; littorinimorpha and neogastropoda are clades within the hypsogastropoda. the clade opisthobranchia contains the clade cephalaspidea. changes in taxonomy since this classification have also been considered, e.g. williams et al. (2008), williams et al. (2010), bandel (2006), fehse (2007) and geiger (2012). worms editorial board (2014), world register of marine species, has been consulted. the frequencies of the species are indicated: abundant: >500 specimens; very common: 100–499 specimens; common: 50–99 specimens; rather common 20–49 specimens; rare: 5–19 specimens and very rare: 1–4 specimens. for rare and very rare species, the depository and numbers of specimens are given. 3636 phylum mollusca linnaeus 1758 class gastropoda cuvier 1795 clade vetigastropoda superfamily fissurellioidea fleming 1822 family fissurellidae fleming 1822 subfamily fissurellinae fleming 1822 genus rimula lowe 1852 type species. rimula blainvillei defrance 1827. rimula? sp. (fig. 1) additional material. arf, 8 specimens (øsm-1005925116); isl, 9 specimens; mno, 3 specimens. the species is rather common. remarks. the species is assigned to rimula with some doubt, as this species generally has a long and narrow perforation. furthermore, no specimen allows examination of the internal characters of the shell to ascertain whether or not there is an internal septum. fig. 1. rimula? sp. , a, b: mguh 30963 (ex rkz 154) 1.6 x 2.5 mm, 2.5 x 1.5 mm. c: mguh 30964 (ex rkz 154a). a b c subfamily emarginulinae children 1834 genus emarginula lamarck 1801 type species. emarginula conica lamarck 1801. emarginula coralliora lundgren 1867 (fig. 2) 1867 emarginula coralliorum (m.u.h.) lundgren, p. 19, plate 1, figs 5a, b. 1902 emarginula coralliorum (m.u.h.) lundgren – ravn, p. 213, plate 1, figs 1, 2. 1933 emarginula coralliorum (m.u.h.) lundgren – ravn, p. 20, plate 1, figs 13a, b. 2004 emarginula (s. s.) coralliora lundgren 1867 – pacaud, p. 599, fig. 10. type material. holotype lo 180t (lundgren 1867, fig. 5a), paratype lo 181t (lundgren 1867, fig. 5b). additional material. gm, 10 specimens (ravn 1933). the species is very common. a b fig. 2. emarginula coralliora lundgren 1867. mguh 3122, height 3.2 mm, length 4.7 mm. superfamily pleuroromarioidea swainson 1840 family pleurotomariidae swainson 1840 genus leptomaria eudes-deslongchamps 1864 type species. pleurotomaria amoena eudes-deslongchamps 1849. leptomaria niloticiformis (von schlotheim 1820) (fig. 3) 1820 trochilites niloticiformis schlotheim, p. 156. 1902 pleurotomaria niloticiformis (schlotheim) – ravn, p. 214, plate 1, fig. 3, fig. 4? 1933 pleurotomaria niloticiformis (schlotheim) – ravn, p. 25, plate 1, fig. 10. additional material. mguh 3119. small specimens are very common, whereas larger complete specimens are rare. the species is very common. remarks. ravn (1902b) described the species on the basis of two specimens with the shell preserved and discussed the variability of the species. pacaud (2004, p. 619) discussed the validity of the taxon pleurotomaria niloticiformis von schlotheim 1820. this species was compared with trochilites politus and considered as a new species because of its large size (von schlotheim 1820). however, neither descriptions nor illustrations of p. niloticiformis were given by von schlotheim (1820). 37 ravn (1902b, 1933) and pacaud (2004) suggested the possibility of several species of pleurotomaria at faxe. pacaud described leptomaria meyeri from the danian of the paris basin and stated that this species was also present in the danian at faxe. the danish material from the baunekule facies differs from the french specimens by having a smaller apical angle (approx. 70° instead of 90°–120°) and almost flat whorls with very weak spiral ornamentation. further studies of the large leptomaria material from faxe are necessary, but leptomaria meyeri is probably present. a c b d fig. 3. leptomaria niloticiformis (von schlotheim 1820). a–c: mguh 30965 (ex rkz 105), width 1.5 mm, height 1.1 mm. d: mguh 30966 (ex isl), height 23.0 mm, width 26.2 mm. leptomaria sp. (fig. 4) material. only the illustrated specimen is known. description. the specimen consists of the protoconch and 11/4 teleoconch. the nucleus is larger than on the previous species and the juvenile shell has a stronger sculpture. remarks. ravn (1933) noted the possible presence of several species of leptomaria. pacaud (2004) described the species leptomaria penultima (d’orbigny 1850), which is also known from the danian at limhamn, sweden. an external impression in arf shows a similar sculpture and it seems possible that the illustrated juvenile specimen might belong to this species. however, further studies are necessary. fig. 4. leptomaria sp. mguh 30967 (ex rkz 106), maximum diameter 2.2 mm. superfamily haliotoidea rafinesque 1815 family temnotropidae cox 1960 genus temnotropis laube 1870 type species. temnotropis carinata (münster 1841). temnotropis sp. (fig. 5) additional material. gm, 3 specimens (gm 1991.4390, gm 1991.4415, gm 1991.4221); isl, 1 specimen; mno, 1 specimen. the species is rare. a b fig. 5. temnotropis sp. mguh 30968 (ex rkz 112), maximum diameter 1.9 mm. superfamily porcellioidea koken 1895 family porcellidae koken 1895 genus faxetrochus schnetler & lozouet 2012 type species. faxetrochus problematicus schnetler & lozouet 2012. faxetrochus problematicus schnetler & lozouet 2012 (fig. 6) 2012 faxetrochus problematicus schnetler & lozouet, p. 4, plate 1a–f. 3838 material. only the holotype (mguh 29810) is known. fig. 6. faxetrochus problematicus schnetler & lozouet 2012. mguh 29810, height 1.8 mm, width 2.8 mm. photo: pierre lozouet, paris. superfamily scissurellioidea gray 1847a family scissurellidae gray 1847a subfamily scissurellinae gray 1847a genus scissurella d’orbigny 1824 type species. scissurella laevigata d’orbigny 1824. scissurella annulata ravn 1933 (fig. 7) 1933 scissurella annulata ravn, p. 25, plate 1, figs 14a–c. 1998 maxwellella annulata (ravn 1933) – bandel, p. 19 ; plate 5, figs 4–6. 2001 maxwellella annulata (ravn 1933) – schnetler et al., p. 86, plate 1, figs 7–9; plate 3, figs 4a–c. type material. holotype mguh 3123. additional material. mguh 25753; gm 3495, 24 specimens; gm 1991.3499, 1 specimen; gm 1991.4415, 1 specimen; mnhn, 2 specimens; jmp, 2 specimens; smf, 2 specimens (smf 321222); isl, 22 specimens; arf, 54 specimens; mno 19 specimens. the species is very common. remarks. according to geiger (2012), maxwellella bandel 1998 is a junior synonym of scissurella d’orbigny 1824. fig. 7. scissurella annulata ravn 1933. mguh 3123, height 0.8 mm, width 1.5 mm. scissurella (s. l.) aliceae schnetler, lozouet & pacaud 2001 (fig. 8) 2001 scissurella aliceaea schnetler et al., p. 82, plate 1, figs 1–3; plate 3, figs 1, 2. type material. holotype mguh 25748; paratypes mguh 25749, mguh 25750, mguh 25758, mguh 25759. additional material. gm 1991.3496, 6 specimens; gm 1991.4415, 2 specimens; mnhn, 1 specimen (mnhnlp r63044); smf, 1 specimen (smf 321220); isl, 7 specimens; arf, 49 specimens (øsm-10059-21734, 25848 and 25851); mno, 2 specimens. the species is common, but easily overlooked. 39 fig. 8. scissurella (s. l.) aliceae schnetler, lozouet & pacaud 2001. a, b: mguh 25759 (ex rkz 119), width 1.9 mm, height 1.6 mm. c, d: mguh 25758 (ex rkz 120), width 1.8 mm. e: mguh 30969 (ex rkz 100), height 1.8 mm. scissurella ravni (schnetler, lozouet & pacaud 2001) (fig. 9) 2001 praescissurella? ravni schnetler et al., p. 84, plate 1, figs 4–6; plate 3, figs 5a–c. type material. holotype mguh 25752. paratypes mguh 25751, mguh 25761 (ex rkz 121). additional material. gm 1991.3497, 4 specimens; mnhn, 1 specimen; jmp, 1 specimen (mnhnlpr63045); smf, 1 specimen (smf 321221); isl, 3 specimens; arf, 1 specimen (øsm-10059-25843). the species is rare. remarks. according to geiger (2012), praescissurella lozouet 1998 is a junior subjective synonym of scissurella d’orbigny 1824. fig. 9. scissurella ravni (schnetler, lozouet & pacaud 2001). mguh 25761 (ex rkz 121), height 1.2 mm, width 1.6 mm. family anatomidae mclean 1989 genus anatoma woodward 1859 type species. scissurella crispata fleming 1828. anatoma hedegaardi (bandel 1998) (fig. 10) 1998 scissurella hedegaardi n. sp. – bandel, p. 16, plate 4, figs 3, 4. 2001 anatoma hedegaardi (bandel 1998) – schnetler et al., p. 86, plate 2, figs 1–3. type material. sgpih nr. 3872 additional material. mguh 25755; gm 1991.4221 and gm1991.4222, 2 specimens (leg. g.v. olsen); arf, 3 specimens (øsm-10059-25850); isl, 1 specimen. the species is rare. fig. 10. anatoma hedegaardi (bandel 1998). mguh 25755, height 2.1 mm. 4040 anatoma rosenkrantzi schnetler, lozouet & pacaud 2001 (fig. 11) 2001 anatoma rosenkrantzi schnetler et al., p. 8, plate 2, figs 4–6; plate 3, figs 3a–c. type material. holotype mguh 25756. paratypes mguh 25757 and mguh 25760. material. no further specimens are known. the species is very rare. fig. 11. anatoma rosenkrantzi schnetler, lozouet & pacaud 2001. a–c: mguh 25760 (ex rkz 118), width 1.5 mm, height 1.1 mm. superfamily trochoidea rafinesque 1815 family trochidae rafinesque 1815 subfamily monodontinae gray 1857 genus osilinus philippi 1847 type species. trochus turbinatus born 1778. osilinus carinatus (ravn 1933) (fig. 12) 1933 monodonta (osilinus) carinata ravn, p. 31, plate 2, figs 6a–d, 7a–d. type material. holotype mguh 3131, paratype mguh 3130. material. the species is abundant. fig. 12. osilinus carinatus (ravn 1933). a–c: mguh 3130, height 1.4 mm, width 1.8 mm. d, e: mguh 3131, height 2.5 mm, width 2.5 mm. genus jujubinus monterosato 1884 type species. trochus matoni payraudeau 1826 jujubinus sp. (fig. 13) material. only the illustrated specimen is known (ex isl, leg. alice rasmussen). remarks. the species resembles jujubinus hannonicus (rutot in cossmann 1915), illustrated by glibert (1973, plate 1, fig. 18). 41 fig. 13. jujubinus sp. mguh 30970 (ex isl), height 2.3 mm, width 2.1 mm. subfamily stomatellinae gray 1840 genus stomatella lamarck 1816 type species. stomatella auricula lamarck 1816. stomatella sp. (fig. 14) material. only the illustrated specimen is known. remarks. the species has slightly convex whorls, a relatively high spire and no umbilicus. the aperture is elongate-ovate. the genus stomatolina (iredale 1937) has a higher spire. species of the genus stomatia (helbling 1779) also have a higher spire and are often strongly sculptured. rosenkrantz suggested, based on the drawing, that it was a natica sp. a b c d fig. 14. stomatella sp. mguh 30971 (ex rkz 110), height 10.2 mm, width 13.7 mm. family turbinidae rafinesque 1815 subfamily skeneinae clark 1851 genus leucorhynchia crosse 1867 type species. leucorhynchia caledonica crosse 1867. leucorhynchia marginata ravn 1933 (fig. 15) 1933 leucorhynchia marginata ravn, p. 27, plate 2, figs 1a–c. type material. holotype mguh 3125. additional material. mno, 2 juvenile specimens; isl, 1 juvenile specimen. the holotype is the only known adult specimen. the species is very rare. remarks. the juvenile specimens have a wide umbilicus and fine spirals on the base. under the adapical suture, the shell has the same spiral as the adult specimen. http://en.wikipedia.org/wiki/skeneinae 4242 fig. 15. leucorhynchia marginata ravn 1933. a–d: mguh 3125, height 4.0 mm, width 5.0 mm. e–g: juvenile specimen mguh 30972 (ex rkz 102), height 0.95 mm, width 1.4 mm. genus skeneoides warén 1992 type species. delphinula exilissima philippi 1844 skeneoides sp. (fig. 16) additional material. arf, 1 specimen; ecs, 1 specimen (øsm 10061f 2-20); mno, 4 specimens; isl, 1 specimen. the species is rare. remarks. the shell is very small and discoidal with a maximum diameter of 0.8 mm. the nucleus is relatively large and the two teleoconch whorls are quickly and regularly increasing in strength. the last whorl is higher than the other whorls and in umbilical view all whorls are visible. the axial sculpture consists of 15 strong orthocline ribs, which are old apertures. the holostomate aperture is subcircular. fig. 16. skeneoides sp. mguh 30973 (ex mno), height 0.45 mm, diameter 0.8 mm. family ataphridae cossmann 1915 genus ataphrus gabb 1869 type species. ataphrus crassus gabb 1869 ataphrus sp. (fig. 17) material. only the illustrated specimen is known. the species is very rare. remarks. the species has convex whorls, a narrow, but deep umbilicus and a subcircular aperture. fig. 17. ataphrus sp. mguh 30974 (ex rkz 141), height 3.5 mm, width 5.0 mm. http://en.wikipedia.org/w/index.php?title=ataphrus&action=edit&redlink=1 43 subfamily margaritinae thiele 1924 genus margarites gray 1847a type species. helix margarita montagu 1808. margarites bruennichi (ravn 1933) (fig. 18) 1933 eumargarita brünnichi ravn, p. 32, plate 3, figs 11a–c. type material. holotype mguh 3144. additional material. the species is abundant. remarks. according to williams et al. (2008), margaritinae thiele 1924 has been moved to turbinidae rafinesque 1815. fig. 18. margarites bruennichi (ravn 1933). mguh 3144, height 2.2 mm, width 2.9 mm. superfamily angarioidea williams et al. 2008 family angariidae gray 1857 subfamily angariinae gray 1857 genus angaria röding 1798 type species. turbo delphinus linnaeus 1758. angaria depressa (ravn 1933) (fig. 19) 1933 delphinula depressa ravn, p. 29, plate 2, figs 3a–c. type material. holotype mguh 3127. material. the species is very common. in a few cases the operculum is found in situ; isolated opercula are rare. fig. 19. angaria depressa. a–d: mguh 30975 (ex rkz 96), height 4.3 mm, width 5.4 mm. e: specimen with operculum in situ mguh 30976 (ex rkz 122), height 2.2 mm, width 3.2 mm. superfamily phasianelloidea williams et al. 2008 family colloniidae cossmann in cossmann & peyrot 1917 subfamily colloniinae cossmann in cossmann & peyrot 1917 genus collonia j.e. gray 1850 type species. collonia marginata (lamarck 1804). collonia (circulopsis) pusilla ravn 1933 (fig. 20) 1933 collonia pusilla ravn, p. 27, plate 1, figs 15a–c. type material. holotype mguh 3124. additional material. the species is very common. a single specimen with the operculum in situ has been found in arf (øsm-10059-25130). 4444 fig. 20. collonia (circulopsis) pusilla ravn 1933. mguh 3124, height 2.0 mm, width 2.9 mm. genus vexinia cossmann 1918 type species. vexinia crassa (baudon 1853). vexinia sp. (fig. 21) additional material. isl, 3 specimens. the species is very rare. fig. 21. vexinia sp. mguh 30977 (ex rkz 111), height 4.3 mm, width 6.1 mm. superfamily eucycloidea koken 1897 family eucycloscalidae gründel 2007 genus eucycloscala cossmann 1895 type species. turbo davoustii d’orbigny 1850. eucycloscala ultima ravn 1933 (fig. 22) 1933 eucycloscala ultima ravn, p. 28, plate 2, figs 2a, b. 2010 eucycloscala ultima ravn 1933 – bandel, p. 440, figs 2 h, i. type material. holotype mguh 3126. material. the species is rather common. a b fig. 22. eucycloscala ultima ravn 1933. mguh 3126, height 2.3 mm, width 2.0 mm. superfamily seguenzioidea verril 1884 family chilodontidae wenz 1938 genus danilia brusina 1865 type species. danilia otaviana cantraine 1835. danilia faxensis (ravn 1933) (fig. 23) 1933 monodonta (danilia) faxensis ravn, p. 29, plate 2, figs 4a–c. 2010 danilia faxensis (ravn 1933) – bandel, p. 468, figs 12e–g. type material. holotype mguh 3128. material. the species is abundant. remarks. bandel (2010) referred the genus danilia to the family turcicidae bandel 2010. bouchet & gofas (2014) referred it to the family chilodontidae wenz 1938. http://en.wikipedia.org/wiki/seguenzioidea 45 a b fig. 23. danilia faxensis (ravn 1933). mguh 30978 (ex isl), height 4.5 mm, width 3.2 mm. danilia quadricordata (ravn 1933) (fig. 24) 1933 monodonta (danilia) quadricordata ravn, p. 30, plate 2, figs 5a–c. 2010 danilia quadricordata (ravn 1933) – bandel, p. 468, figs 12a–d. type material. mguh 3129. material. the species is abundant. fig. 24. danilia quadricordata (ravn 1933). mguh 3129, height 5.0 mm, width 3.9 mm. danilia fenestrata (ravn 1933) (fig. 25) 1933 monodonta (danilia) fenestrata ravn, p. 31, plate 2, figs 8a–c. 2010 danilia fenestrata (ravn 1933) – bandel, p. 468, figs 12h–j. type material. mguh 3132. material. the species is abundant. fig. 25. danilia fenestrata (ravn 1933). mguh 3132, height 3.2 mm, width 2.4 mm. subfamily chilodontinae wenz 1938 chilodontinae gen. et sp. indet. (fig. 26) additional material. gm 1991.4398, 1 specimen. the species is very rare. remarks. rosenkrantz suggested the genus name kangilia in his files, but never published it. the species resembles chilodontinae, new genus, species 2 from the paleocene of nuussuaq, west geenland, illustrated by kollmann & peel (1983, p. 25, figs 23a, b). most likely rosenkrantz proposed the name because of the similarity of the two species. the faxe species differs by having a much finer sculpture and a base demarcated by an angulation. in general outline and aperture it resembles the species from nuussuaq. only the illustrated specimen (fig. 26b) and one additional specimen (gm 1991.4398) have a rather well preserved aperture. two specimens are known from the coral limestone in faxe (rkz 99, isl). a b c d fig. 26. chilodontinae gen. et sp. indet. a–c: mguh 30980 (ex isl), height 4.5 mm, width 5.4 mm. d: mguh 30979 (ex rkz 98), height 10.0 mm, width 11.2 mm. 4646 clade caenogastropoda cox 1960 clade sorbeoconcha ponder & lindbergh 1997 superfamily cerithioidea fleming 1822 family cerithiidae fleming 1822 genus ataxocerithium tate 1894 type species. cerithium serotinum a. adams 1855. ataxocerithium faxensis (ravn 1933) (fig. 27) 1933 cerithiopsis faxensis ravn, p. 50, plate 5, figs 6a, b, 7a, b, 8a, b. 1998 ataxocerithium faxensis (ravn 1933) – nützel, p. 116, plate 16, figs a, b. type material. holotype mguh 3169, paratype mguh 3170. additional material. gm (ravn), 119 specimens; smf 311 741, 1 specimen. the species is very common. remarks. neither ravn (1933) nor nützel (1998) found spiral ornament on the protoconch and on the drawing (fig. 27) only axial riblets are seen. however, the terminal half protoconch whorl has two very fine spiral riblets above the abapical suture and they increase in strength towards the transition into the teleoconch. thus the protoconch matches the protoconch of the genus ataxocerithium. a b fig. 27. ataxocerithium faxensis (ravn 1933). a: mguh 3169, height 5.2 mm, width 1.6 mm. b: mguh 3170, height 2.7 mm, width 1.2 mm. ataxocerithium sp. 1 (fig. 28) additional material. arf, 16 specimens (øsm-1005925046); gm 1991.4279, 1 badly preserved specimen; isl, 1 specimen. the species is rare. remarks. the species differs from the two preceding species by having opisthocline axial ribs, which are stronger than the spirals. the protoconch has axial riblets. fig. 28. ataxocerithium sp. 1. øsm-10059-25046 (ex arf), height 3.9 mm, width 1.1 mm. ataxocerithium sp. 2 (fig. 29) additional material. gm 1991.4204, 1 specimen; gm 1991.4305, 3 specimens. the species is very rare. remarks. the protoconch has weak axial ribs and abapically two spiral riblets. the teleoconch whorls have an adapical spiral with strong knobs, a spiral with weak knobs on the middle of the whorl and two abapical spiral furrows. the number of knobs is c. 15. 47 fig. 29. ataxocerithium sp. 2. mguh 30982 (ex gm 1991.4205), height 4.1 mm, width 1.4 mm. genus bittium leach in gray 1847b type species. bittium reticulatum da costa 1778. bittium cf. transenna (bayan 1873) (fig. 30) 1933 cf. bittium transenna (bayan 1873) – ravn, p. 44, plate 3, figs 14a, b. additional material. arf, 21 specimens (øsm-1005925027, 25058, 25059, 25063, 25064, 25094); isl, 11 specimens. the species is rather common. remarks. ravn (1933) had only one incomplete specimen at hand. better material is now at our disposal, but comparison with french material is necessary for a secure assignment to the species of bayan (1873). fig. 30. bittium cf. transenna (bayan 1873). mguh 3147, height 4.2 mm, width 1.3 mm. family metacerithiidae cossmann 1906 genus metacerithium cossmann 1906 type species. cerithium trimonile michelin 1838. metacerithium? sp. (fig. 31) material. only the illustrated specimen is known. remarks. the species has almost smooth whorls and the aperture is not preserved. however, whorls are characteristic for the genus metacerithium. fig. 31. metacerithium? sp. mguh 30983 (ex rkz 195), height 4.7 mm. 4848 superfamily campaniloidea douvillé 1904 family campanilidae douvillé 1904 genus campanile bayle in fischer 1864 type species. cerithium giganteum lamarck 1804. campanile? pseudotelescopium (ravn 1902b) (fig. 32) 1902 cerithium (campanile?) pseudotelescopium (m.u.h.) ravn, p. 219, plate 1, figs 13, 14. 1933 cerithium (campanile?) pseudotelescopium (m.u.h.) ravn, – ravn, p. 45 (partim), non plate 4, figs 8a, b. type material. syntypes mguh 87 and mguh 88. material. most specimens are incomplete. the species is rather common. discussion. ravn (1933, p. 45) referred the species to campanile with a query, since no specimens with a well preserved aperture and canal were available. such specimens have not been found as yet. ravn referred with some doubt the illustrated juvenile specimen named below as campanile? sp. 1 to this species. this specimen has about 18 axial ribs, while campanile? pseudotelescopium has about 30. a b fig. 32. campanile? pseudotelescopium (ravn 1902b). mguh 30984 (ex rkz 181), height 28.0 mm, width 9.7 mm. campanile? subglabra (ravn 1933) (fig. 33) 1933 newtoniella subglabra ravn, p. 47, plate 5, figs 5a, b. type material. holotype mguh 3167. material. the species is rare and only represented by fragments. remarks. ravn referred the species to the genus newtonilla, based on a single fragmentary specimen. the whorls are flat and relatively low, with three fine spirals on the adapical half of the whorl. the species has some resemblance with the genus campanile, but the assignment is questionable, as only fragments are available. the protoconch is unknown. fig. 33. campanile? subglabra (ravn 1933). mguh 3167, height 6.8 mm, width 2.3 mm. campanile? sp. 1 (fig. 34) 1933 cerithium (campanile?) pseudotelescopium (m.u.h.) ravn, – ravn, p. 45 (partim), plate 4, figs 8a, b (non ravn). material. only the illustrated specimen is known. 49 remarks. the species has been discussed above. the drawing in the rosenkrantz files is not correct, as the shell is drawn too slender. fig. 34. campanile? sp. 1. mguh 3155, height 7.2 mm, width 2.2 mm. campanile? sp. 2 (fig. 35) additional material. gm, 1 specimen (1991.4128). remarks. the species has a subsutural spiral with c. 15 knobs on the first teleoconch whorls. on the later whorls there are two close-set adapical fine spirals under the suture and a spiral furrow a little above the middle of the whorl. the whorls are slightly convex and relatively low and the aperture is rather narrow and rounded rectangular. in general outline and sculpture, the species has some resemblance with campanile? subglabra. fig. 35. campanile? sp. 2. mguh 30985 (ex rkz 70), height 12.3 mm, width 3.0 mm. family trypanaxidae gougerot & le renard 1987 genus trypanaxis cossmann 1889 type species. trypanaxis umbilicata (lamarck 1804). trypanaxis faxensis ravn 1933 (fig. 36) 1933 trypanaxis? faxensis ravn, p. 53, plate 5, figs 18a, b. type material. mguh 3280. material. the species is rather common. 5050 a b c fig. 36. trypanaxis faxensis ravn 1933. mguh 30986 (ex rkz 67), height 7.2 mm, width 3.0 mm, height of protoconch 2.0 mm. family turritellidae lovén 1847 genus mesalia gray 1847a type species. cerithium meaal adanson 1757 mesalia? sp. 1 (fig. 37) additional material. gm, 7 specimens (1991.4092 1991.4094 and 1991.4106); isl, 7 specimens; mno, 1 specimen. the species is rare. remarks. this and the following two species have a spiral ornament, which resembles the ornament of the mesalia or sigmesalia species. however, they differ from these genera by having a rather well-developed canal. thus, the assignment to the genus mesalia of this and the next two species is questionable. rosenkrantz assigned these species to the genus orthochetus cossmann 1889 in his notes, but this genus has a completely different sculpture. mesalia? sp. 1 is characterised by a rather slender outline and three equidistant spirals of equal strength. fig. 37. mesalia? sp. 1. mguh 30987 (ex rkz 71), height 8.1 mm, width 2.8 mm. mesalia? sp. 2 (fig. 38) additional material. gm, 2 specimens (1991.4091 and 1991.4093); isl, 4 specimens; ebp, 1 specimen. the species is rare. remarks. this species is less slender and has two spirals on the abapical part of the whorls. fig. 38. mesalia? sp. 2. mguh 30988 (ex rkz 145), height 2.3 mm, width 1.4 mm. 51 mesalia? sp. 3 (fig. 39) additional material. gm, 1 specimen (1991. 4095); isl, 1 specimen. the species is very rare. remarks. this species has three spirals of equal strength, but the shell is less slender than mesalia? sp. 1 and the whorls are more convex. fig. 39. mesalia? sp. 3. mguh 30989 (ex rkz 187), height 3.2 mm, width 1.4 mm. family siliquariidae anton 1838 genus tenagodus guettard 1770 type species. serpula anguina linnaeus 1758. tenagodus ornatus (lundgren 1867) (fig. 40) 1867 siliquaria ornata (m. u. h.) lundgren, p. 17, plate 1, fig. 4. 1902 siliquaria ornata (m. u. h.), lundgren – ravn, p. 219; plate 1, figs 11, 12. 1933 siliquaria ornata (m. u. h.), lundgren – ravn, p. 40. material. the species is rare and represented by juvenile specimens. fig. 40. tenagodus ornatus (lundgren 1867). mguh 30990 (ex rkz 107), maximum diameter 2.2 mm. tenagodus sp. (fig. 41) material. only the illustrated specimen is known. it differs from the preceeding species by not having disjunct whorls. fig. 41. tenagodus sp. mguh 30991 (ex rkz 144), height 5.9 mm, width 5.5 mm. 5252 clade hypsogastropoda ponder & lindberg 1997 clade littorinimorpha golikov & starobogatov 1975 superfamily capuloidea fleming 1822 family capulidae fleming 1822 genus trichotropis broderip & sowerby 1829 type species. turbo bicarinatus sowerby 1825. trichotropis? sp. (fig. 42) material. only the illustrated specimen is known. description. the species has a protoconch with c. 2½ whorls, of which the last two have prosocline axial ribs. the teleoconch whorls are medium convex with a flat adapical ramp and separated by a deep suture. they have three sharp primary spirals, separated by deep furrows. four additional spirals appear on the abapical part of the whorl, decreasing. the base has a similar spiral ornament. the axial ribs are only visible on the terminal half whorl, where they cause weak knobs on the primary spirals. the aperture is ovate and has a weak spout near the columella. discussion. the protoconch and teleoconch features suggest that the species could be referred to the genus trichotropis (s. kiel, personal communication 2014). kollmann & peel (1983, p. 59, fig. 116) illustrated a species with a similar spiral ornament as trichotropinae, new genus. a b fig. 42. trichotropis? sp. øsm-10059-25092 (ex arf), height 3.7 mm, width 2.4 mm. capulidae gen. et sp. indet. (fig. 43) material. only the illustrated specimen is known. a b c fig. 43. capulidae gen. et sp. indet. mguh 30992 (ex rkz 109), height 0.9 mm, width 1.3 mm. superfamily cypraeoidea rafinesque 1815 family cypraeidae rafinesque 1815 subfamily gisortiinae schilder 1927 genus palaeocypraea schilder 1928 type species. cypraeacites spiratus von schlotheim 1820. palaeocypraea spirata (von schlotheim 1820) (fig. 44) 1902 cypraea spirata schloth. – ravn, p. 21, plate 2, fig. 2. 1928 palaeocypraea spirata schloth. – schilder, p. 19, figs 3–5, 12–14. 1933 cypraea (palaeocypraea) spirata schloth. – ravn, p. 57, plate 6, figs 4a,b, 5a, b. type material. mguh 3184, mguh 3185. material. juvenile specimens are common, but complete adult specimens are rather common. remarks. ravn (1933, p. 57, plate 6, figs 4a, b, 5a, b) described and illustrated juvenile specimens with a well preserved protoconch and he also mentioned intermediate specimens with protoconch and teleoconch whorls preserved. on adult specimens the protoconch is completely hidden. 53 a b c d fig. 44. palaeocypraea spirata (von schlotheim 1820). a, b: mguh 30994 (ex rkz 33), height 16.7 mm, width 11.2 mm. c, d: mguh 3185, height 4.7 mm, width 2.8 mm. palaeocypraea sp. (fig. 45) additional material. rkz 32. the species is very rare. remarks. only two juvenile specimens are known. they differ from palaeocypraea spirata by having a smaller protoconch, a coarser diagonal cancellation of the terminal protoconch whorls and especially by the spiral ornament and axial sculpture of the first teleoconch whorl. this species has two coarse spirals on the whorl and weaker spirals adapically and abapically. in between the strong spirals there are three much weaker spirals. the axial ribs cause knobs on the two strong spirals, about 20 on a whorl. juvenile palaeocypraea spirata have a larger protoconch with a wider apical angle and a finer sculpture. the spirals are finer and more numerous and the spirals have no knobs. the axial sculpture is much finer. rosenkrantz in his notes suggested the name palaeocypraea poulseni for the present species, but he never published it. a b fig. 45. palaeocypraea sp. mguh 30995 (ex rkz 166), height 2.0 mm, width 1.3 mm. genus bernaya jousseaume 1884 type species. cypraea media deshayes 1835. subgenus bernaya (protocypraea) schilder 1927 type species. eocypraea orbignyana vredenburg 1920. bernaya (protocypraea) globuliformis (ravn 1902b) (fig. 46) 1902 cypraea (protocypraea) globuliformis ravn, p. 23, plate 2, fig. 6. 1928 protocypraea globuliformis ravn – schilder, p. 23, fig. 15. 1933 cypraea (protocypraea) globuliformis ravn – ravn, p. 58. type material. mguh 103. additional material. arf, 1 specimen (øsm-1005921043). the species is very rare. 5454 fig. 46. bernaya (protocypraea) globuliformis (ravn 1902b). mguh 30996 (ex rkz 264), height 9.8 mm, width 7.8 mm. a b family eocypraeidae schilder 1924 genus eocypraea cossmann 1903 type species. cypraea inflata lamarck 1802. eocypraea danica (schilder 1928) (fig. 47) 1902 cypraea bullaria schloth. – ravn, p. 22, plate 2, figs 4a–c (non schlotheim). 1928 eocypraea danica schilder, p. 12, figs 1, 2, 11. 1933 cypraea (eocypraea) danica schilder – ravn, p. 58. additional material. arf, 21 specimens (øsm-1005921041). the species is rather common. a b fig. 47. eocypraea danica (schilder 1928). mguh 30997 (ex rkz 34), height 18.4 mm, width 13.4 mm. superfamily littorinoidea children 1834 family littorinidae children 1834 subfamily littorininae children 1834 genus littoraria gray 1833 type species. littorina pulchra sowerby 1832. subgenus littoraria (littorinopsis) mörch 1876 type species. littorina subangulata lamarck (error for littorina angulifera (lamarck 1822)). littoraria (littorinopsis) faxensis (ravn 1933) (fig. 48) 1933 littorinopsis faxensis ravn, p. 35, plate 3, figs 3a, b. type material. holotype mguh 3136. additional material. gm, 3 specimens; mno, 1 specimen. the species is rare. a b c fig. 48. littoraria (littorinopsis) faxensis (ravn 1933). mguh 3136, height 2.2 mm, width 1.6 mm. family pickworthiidae iredale 1917 subfamily pickworthiinae iredale 1917 genus sansonia jousseaume 1892 type species. sansonia tuberculata watson 1886. 55 sansonia hedegaardi bandel & kowalke 1997 (fig. 49) 1997 sansonia hedegaardi bandel & kowalke, p. 14, plate 5, figs 2, 5. type material. holotype gpium 3768, paratype gpium 3769. material. the species is common, but easily overlooked. fig. 49. sansonia hedegaardi bandel & kowalke 1997. mguh 30998 (ex rkz 160), height 1.5 mm, width 1.0 mm. a b c sansonia sp. (fig. 50) additional material. isl, 2 specimens; mno, 1 specimen. the species is rare and easily overlooked. it is a little more slender than the preceding species and differs by having a prosocline labrum in lateral view. fig. 50. sansonia sp. mguh 30999 (ex rkz 161), height 1.7 mm, width 1.1 mm. genus mareleptopoma moolenbeek & faber 1984 type species. mareleptopoma karpatensis moolenbeek & faber 1984. mareleptopoma? sp. (fig. 51) additional material. arf, 3 specimens (øsm-1005925756); isl, 2 specimens. the species is rare and easily overlooked. it is less slender than the two preceding species. fig. 51. mareleptopoma? sp. mguh 31000 (ex rkz 162), height 1.3 mm, width 1.0 mm. genus urceolabrum wade 1916 type species. urceolabrum tuberculatum wade 1916. urceolabrum sp. 1 (fig. 52) additional material. arf, 1 specimen (øsm-1005925846); isl, 2 specimens. the species is very rare. remarks. the species has rather strong axial ribs on all whorls. 5656 fig. 52. urceolabrum sp. 1. mguh 31001 (ex rkz 114), height 1.5 mm, width 1.8 mm. urceolabrum sp. 2 (fig. 53) material. only the illustrated specimen is known. remarks. this species differs from urceolabrum sp. 1 in that the axial ribs fade out on the terminal whorls and furthermore by possessing a lower apex. fig. 53. urceolabrum sp. 2. mguh 31002 (ex rkz 104), height 1.3 mm, width 2.2 mm. subfamily sherboniinae iredale 1917 genus faxia ravn 1933 type species. faxia macrostoma ravn 1933. faxia macrostoma ravn 1933 (fig. 54) 1933 faxia macrostoma ravn 1933, p. 49, plate 6, figs 9a, b, 10a, b, 11. type material. holotype mguh 3189, paratypes mguh 3190, 3191. additional material. rkz 178. adult specimens with a complete aperture are rare, whereas juvenile specimens are common. fig. 54. faxia macrostoma ravn 1933. a–c: mguh 31003 (ex rkz 178), height 5.4 mm, width 2.6 mm. d–f: mguh 31004 (ex rkz 180), height 1.2 mm, width 0.8 mm, diameter of protoconch 0.6 mm. superfamily naticoidea guilding 1834 family naticidae guilding 1834 naticidae gen. et sp. indet. (fig. 55) material. only the illustrated specimen is known. 57 remarks. the species has a low spire, a straight columella and an almost semicircular aperture. there is a narrow umbilicus. as naticids are members of the infauna, it is not surprising that this genus is rare in the baunekule fauna. fig. 55. naticidae gen. et sp. indet. mguh 31005 (ex rkz 103), height 1.7 mm, width 1.9 mm. superfamily rissooidea gray 1847a family rissoidae gray 1847a subfamily rissoininae gray 1847a genus zebina h. & a. adams 1854 type species. zebina browniana d’orbigny 1842. zebina sp. 1 (fig. 56) additional material. arf, one specimen (øsm-1005925080); isl, 10 specimens; mno, 2 specimens. the species is rare. fig. 56. zebina sp. 1. mguh 31006 (ex rkz 143), height 3.5 mm, width 1.5 mm. zebina sp. 2 (fig. 57) additional material. rkz 163b. remarks. the species differs from zebina sp. 1 in having flatter whorls and a more thickened columella. fig. 57. zebina sp. 2. mguh 31007 (ex rkz 163d), height 3.7 mm, width 1.4 mm, diameter of protoconch 0.3 mm. genus pseudotaphrus cossmann 1888 type species. bulimus buccinalis lamarck 1804. pseudotaphrus sp. (fig. 58) material. only the illustrated specimen is known. fig. 58. pseudotaphrus sp. mguh 31013 (ex rkz 132), height 2.7 mm, width 1.6 mm. 5858 rissoininae gen. et sp. indet. (fig. 59) material. the species is rather common. remarks. rosenkrantz suggested eulimella (belonidium) sp. for this material. however, the protoconch is not heterostrophic as on typical pyramidellidae. a. warén (personal communication 2013) suggested that the species was most likely a rissoininae. fig. 59. rissoininae gen. et sp. indet. a: mguh 31008 (ex rkz 49a), height 4.5 mm, width 1.6 mm. b: mguh 31009 (ex rkz 49b), height 5.1 mm, width 1.2 mm. c: mguh 31010 (ex rkz 49c), height 4.6 mm, width 1.0 mm. d: mguh 31011 (ex rkz 49d), height 4.0 mm, width 0.9 mm. e: mguh 31012 (ex rkz 49e), height 4.5 mm, width 0.9 mm. family caecidae gray 1850 genus caecum fleming 1817 type species. caecum tracheum montagu 1803. caecum sp. (fig. 60) material. only the illustrated specimen is known. fig. 60. caecum sp. mguh 31014 (ex rkz 135), height 2.5 mm. family tornidae sacco 1896 (1884) genus circulus jeffreys 1865 type species. delphinula duminyi requien 1848. circulus sp. 1 (fig. 61) material. only the illustrated specimen is known. the species is very rare. remarks. the species has a keel at the periphery and a narrow umbilicus. on the base, five spirals are visible. fig. 61. circulus sp. 1. mguh 31015 (ex rkz 130), height 0.5 mm, width 1.2 mm. circulus sp. 2 (fig. 62) material. only the illustrated specimen is known. http://en.wikipedia.org/wiki/caecidae 59 remarks. this species has a rather sharp carina on the periphery and a narrow umbilicus and resembles circulus (s. str.) aurelius (d’orbigny, 1850), illustrated as circulus laevigatus (deshayes 1862) by gougerot (1970, p. 39, fig. 7). fig. 62. circulus sp. 2. mguh 31016 (ex rkz 117), height 0.9 mm, width 1.8 mm. subfamily teinostomatinae cossmann in cossmann & peyrot 1917 genus teinostoma h. & a. adams 1853 type species. teinostoma politum a. adams in h. & a. adams 1853. teinostoma glaberrimum ravn 1933 (fig. 63) 1933 tinostoma glaberrimum ravn, p. 33, plate 2, figs 9a–c. type material. holotype mguh 3133. additional material. arf, 4 specimens (øsm-1005921431); isl, 3 specimens; mno, 3 specimens. the species is rare. fig. 63. teinostoma glaberrimum ravn 1933. mguh 3133, height 3.4 mm, width 6.1 mm. genus sigaretornus iredale 1936 type species. adeorbis sigaretinus pilsbry 1897. sigaretornus sp. (fig. 64) material. the illustrated specimens are the only ones known. fig. 64. sigaretornus sp. a–c: mguh 31017 (ex rkz 116), height 0.6 mm, width 1.7 mm. d: mguh 31018 (ex rkz 115), diameter 2.0 mm. 6060 superfamily tonnoidea suter 1913 (1825) family ranellidae gray 1854 genus ranella lamarck 1816 type species. ranella gigantea lamarck 1816. ranella faxensis ravn 1933 (fig. 65) 1933 ranella faxensis ravn, p. 59, plate 6, figs 6a, b. type material. holotype mguh 3186. additional material. gm, 2 specimens; arf, 11 specimens (øsm-10059-25002 and øsm-10059-25819); isl, 1 specimen. the species is rare. fig. 65. ranella faxensis ravn 1933. mguh 3186, height 4.0 mm, width 1.8 mm. subfamily cymatiinae iredale 1913 (1854) genus sassia bellardi 1873 type species. sassia apenninica (sasso 1827). sassia faxense (ravn 1933) (fig. 66) 1902 tritonium fenestratum ravn, p. 227, plate 2, figs 7, 8 (non vincent 1878). 1933 tritonium (sassia) faxense ravn, p. 58, plate 5, figs 16a, b; plate 6, figs 7a, b. type material. syntypes to tritonium fenestratum are mguh 104 and mguh 105. syntypes to sassia faxensis are mguh 3178 and mguh 3187. material. juvenile specimens are rather common, adult specimens are rare. remarks. the species has a sculpture of rather coarse spirals and radial ribs of almost the same strength, resulting in a pattern of quadrates. on the columella are two or three folds and the labrum is thickened and has knobs internally. fig. 66. sassia faxense (ravn 1933). a, b: øsm-10059-21047 (ex arf), height 5.8 mm, width 3.9 mm. c, d: mguh 31020 (ex isl), height 9.9 mm, width 5.2 mm. sassia sp. (fig. 67) additional material. isl, 3 specimens. the species is very rare. remarks. the protoconch is smaller and more slender than sassia faxense and has a sculpture of three fine spiral riblets on the last whorl. on the terminal half whorl, finer secondary spiral riblets are inserted between the 61 three spirals and in between the abapical spire and the abapical suture. the teleoconch whorls have a sculpture of fine spiral ribs with secondary spirals and fine radial ribs. fine knobs occur at the intersections. the canal is relatively long. fig. 67. sassia sp. mguh 31021 (ex isl), height 7.6 mm, width 4.0 mm. genus cymatium röding 1798 type species. murex femorale linnaeus 1758 subgenus cymatium (monoplex) perry 1811 type species. monoplex australasiae perry 1811. cymatium (monoplex) subglabrum (ravn 1902b) (fig. 68) 1902 tritonium subglabrum ravn, p. 228, plate 2, figs 9, 10. 1933 tritonium (lampusia?) subglabrum ravn, p. 59, plate 5, fig. 15. type material. syntypes mguh 106 and mguh 107. material. the species is rather common. protoconchs with a fine spiral ornament and more or less defective larger specimens have been found. fig. 68. cymatium (monoplex) subglabrum (ravn 1902b). mguh 31022, height 54.0 mm, width 30.1 mm. superfamily truncatelloidea gray 1840 family elachisinidae ponder 1985  genus laeviphitus van aartsen, bogi & giusti 1989 type species. laeviphitus verduini van aartsen, bogi & giusti 1989. laeviphitus sp. (fig. 69) material. only the two illustrated specimens are known. 6262 fig. 69. laeviphitus sp. a–c: mguh 31023 (ex rkz 137), height 2.2 mm, width 1.4 mm. d, e: mguh 31024 (ex rkz 133), height 2.2 mm, width 1.4 mm. superfamily vanikoroidea gray 1840 family hipponicidae troschel 1861 genus hipponix defrance 1819 type species. hipponix cornucopia lamarck 1803. hipponix sp. (fig. 70) 1933 hipponyx sp. (i) – ravn, p. 36. additional material. arf, 5 specimens (øsm-1005925823). the species is rare. fig. 70. hipponix sp. mguh 31025 (ex rkz 159), height 0.9 mm, width 1.8, length 2.8 mm. genus eoatlanta cossmann 1889 type species. cyclostoma spiruloides lamarck 1804. eoatlanta ravni schnetler 2013 (fig. 71) 1933 eoatlanta spiruloides (lam.) – ravn, p. 70, plate 7, figs 10a–c, 11 (non lamarck 1804). 2013 eoatlanta ravni schnetler, p. 4, plate 1, figs 1–8. type material. holotype mguh 30376 (leg. s.b. andersen 1972). illustrated paratypes mguh 30377 (leg. s.b. andersen 1972), mguh 30378 (leg. s.b. andersen 1972), rgm 794 231 (leg. s.b. andersen 1972), rgm 794 232 (leg. s.b. andersen 1972), mguh 30379 (ex rkz 108), mguh 30380 (ex rkz 158a), mguh 30381 (ex rkz 158b). additional material. gf 10035-33, 33 specimens; gf 10035-34, 5 specimens; arf, 191 specimens (øsm10059-25124 and øsm-10059-25125). the species is very common. 63 fig. 71. eoatlanta ravni schnetler 2013. a–c: mguh 30379 (ex rkz108), height 0.9 mm, width 2.6 mm. d, e: mguh 30380 (ex rkz 158a), height 0.8 mm, width 2.2 mm. f, g: mguh 30381 (ex rkz 158b), height 0.9 mm, width 2.2 mm. superfamily velutinoidea gray 1840 family triviidae troschel 1863 subfamily eratoinae gill 1871 tribe johnstrupiini schilder 1939 genus johnstrupia ravn 1933 type species. johnstrupia faxensis ravn 1933. johnstrupia faxensis ravn 1933 (fig. 72) 1933 johnstrupia faxensis ravn, p. 61, plate 6, figs 3a, b. type material. mguh 3183. additional material. gm, 2 specimens; isl, 1 specimen. the species is very rare. fig. 72. johnstrupia faxensis ravn 1933. mguh 3183, height 9.6 mm, width 6.4 mm. 6464 informal group ptenoglossa superfamily epitonioidea berry 1910 family epitoniidae berry 1910 genus cerithiscala de boury 1887 type species. cerithiscala primula deshayes 1861. cerithiscala tricincta (ravn 1933) (fig. 73) 1933 tenuiscala (cerithiscala) tricincta ravn, p. 39; plate 3, figs 12a, b. material. the species is rather common. fig. 73. cerithiscala tricincta (ravn 1933) a: mguh 31026 (ex rkz 64a), height 3.6 mm, width 2.3 mm. b: mguh 31027 (ex rkz 64b), height 4.6 mm, width 2.2 mm. c: mguh 31028 (ex rkz 64c), height 3.4 mm, width 1.9 mm. d: mguh 31029 (ex rkz 64d), height 1.8 mm, width 1.0 mm. cerithiscala sp. 1 (fig. 74) additional material. isl, 6 specimens. the species is rare. remarks. the species is considerably more slender than c. tricincta (height/width ratio 3.0 and 1.8 respectively). fig. 74. cerithiscala sp. 1. a: mguh 31030 (ex rkz 63a), height 3.6 mm, width 1.2 mm. b: mguh 31031 (ex rkz 63b), height 4.2 mm, width 1.6 mm. cerithiscala sp. 2 (fig. 75) additional material. gm 1991.4159, 1 specimen; gm 1991.4160, 1 specimen. the species is very rare. description. the species is characterised by having a protoconch, consisting of about seven whorls, and carinated whorls with three primary spirals, running over about 14 axial ribs. the adapical spiral is situated between the adapical suture and the medium spiral which is situated at mid-whorl. the abapical spiral is situated immediately above the abapical suture. the medium spiral causes an angulation of the whorl. the slender shell has a height/ width ratio of 3.0. remarks. the species differs from cerithiscala sp. 1 with respect to its larger protoconch, different spiral ornament and regular convex whorls without a carina. rosenkrantz referred the species to seila on his drawing, but the ornamentation of the teleoconch and the shape of the aperture preclude this assignment. 65 fig. 75. cerithiscala sp. 2. mguh 31032 (ex rkz 68), height 2.3 mm, width 0.8 mm. genus opalia h. & a. adams 1853 type species. scalaria australis lamarck 1822. opalia sp. (fig. 76) additional material. arf, 21 specimens (øsm-1005925068); isl, 5 specimens; mno, 4 specimens. the species is rather common. fig. 76. opalia sp. a, b: mguh 31033 (ex rkz 65a), height 5.1 mm, width 2.6 mm. c: mguh 31034 (ex rkz 65b), height 3.2 mm, width 1.3 mm. genus acrilla h. adams 1860 type species. scalaria acuminata sowerby 1844. acrilla elegans (ravn 1902b) (fig. 77) 1902 scalaria elegans ravn, p. 218, plate 1, fig. 10. 1933 acrilla elegans (ravn) – ravn, p. 37. type material. holotype mguh 84. additional material. 5 specimens (ravn 1902b). the species is very rare. remarks. ravn stated that the specimens were found in the coral limestone. on the drawing of the holotype rosenkrantz has noted “næsekalk” (now referred to as the baunekule facies) and this is in accordance with the state of preservation which is typical of the gastropods from the baunekule facies. 6666 fig. 77. acrilla elegans (ravn 1902b). mguh 84, height 28.0 mm, width 10.0 mm. acrilla sp. 1 (fig. 78) material. only the illustrated specimens are known. remarks. the species has six spirals of almost equal strength and the axial ribs are stronger than the spirals. fig. 78. acrilla sp. 1. a: mguh 31035 (ex rkz 61a), height 3.9 mm, width 2.8 mm. b: mguh 31036 (ex rkz 61b), height 2.4 mm, width 1.1 mm. c: mguh 31037 (ex rkz 61c), height 3.4 mm, width 1.3 mm. d: mguh 31038 (ex rkz 60), height 3.7 mm, width 1.7 mm. acrilla sp. 2 (fig. 79) additional material. mno, 1 specimen. the species is very rare. remarks. the species is more slender than the following species and has three strong spirals and more angular whorls. the axial ribs are stronger than the spirals on the last whorl. 67 fig. 79. acrilla sp. 2. mguh 31039 (ex rkz 59), height 3.3 mm, width 1.1 mm. acrilla sp. 3 (fig. 80) additional material. arf, 1 specimen (øsm-1005925060). the species is very rare. fig. 80. acrilla sp. 3. øsm-10059-25061 (ex arf), height 4.4 mm, width 2.1 mm. acrilla? sp. (fig. 81) additional material. ebp, 1 specimen. the species is very rare. remarks. the species has a slightly coeloconoid outline, convex whorls, which are relatively low. there are c. 10 spirals, which are separated by narrow furrows, and about 12 almost orthocline axial ribs. the aperture is not completely preserved, but it is rather small with a concave columella and a short canal, which is turned to the left. the state of preservation excludes a definitive generic assignment. fig. 81. acrilla? sp. øsm-10059-25105 (ex arf), height 11.5 mm, width 4.4 mm. genus cirsotrema mörch 1852 type species. scalaria varicosa lamarck 1822. cirsotrema sp. (fig. 82) material. only the illustrated specimen is known. 6868 fig. 82. cirsotrema sp. øsm-10059-25042 (ex arf), height 5.1 mm, width 1.7 mm. superfamily eulimoidea philippi 1853 family eulimoidae philippi 1853 genus eulima risso 1826 type species. turbo subulatus donovan 1803. subgenus eulima (polygireulima) cossmann 1894 type species. melania spina grateloup 1838. eulima (polygireulima) danica ravn 1933 (fig. 83) 1933 eulima (polygyreulima) danica ravn, p. 41, plate 4, figs 1a, b, 2a, b. type material. holotype mguh 3149, paratype mguh 3148. material. the species is rather common. fig. 83. eulima (polygireulima) danica ravn 1933. a: mguh 3148, height 5.7 mm, width 2.0 mm. b: mguh 3149, height 3.5 mm, width 1.2 mm. c: mguh 31043 (ex rkz 41a), height 3.7 mm, width 1.7 mm. d: mguh 31044 (ex rkz 41b), height 9.7 mm, width 2.8 mm. eulima (polygireulima) sp. 1 (fig. 84) additional material. arf, 2 specimens (øsm-1005925052 and 25067). the species is very rare. remarks. this species is the most slender of the eulima species. the whorls are slightly to medium convex and the aperture has a spout-like thickening anteriorly. 69 fig. 84. eulima (polygireulima) sp. 1. a: mguh 31045 (ex rkz 42), height 3.0 mm, width 1.0 mm. b: mguh 31046 (ex rkz 43), height 4.9 mm, width 1.3 mm. eulima (polygireulima) sp. 2 (fig. 85) material. only the illustrated specimen is known. remarks. this species has relatively low whorls and a wider aperture, without the anterior spout. fig. 85. eulima (polygireulima) sp. 2. mguh 31047 (ex rkz 44), height 4.8 mm, width 1.6 mm. eulima (polygireulima) sp. 3 (fig. 86) material. only the illustrated specimen is known. remarks. this species has a very fine spiral ornament and a rather narrow aperture with a posterior narrow spout. fig. 86. eulima (polygireulima) sp. 3. mguh 31048 (ex rkz 48), height 4.0 mm, width 1.3 mm. eulima (polygireulima) sp. 4 (fig. 87) material. only the illustrated specimen is known. remarks. this species has a very convex last whorl and a thickened columella (pseudoumbilicus?). the aperture has a posterior narrow spout. 7070 fig. 87. eulima (polygireulima) sp. 4. mguh 31049 (ex rkz 46), height 4.2 mm, width 1.6 mm. subgenus eulima (margineulima) cossmann 1888 type species. eulima fallax deshayes 1862. eulima (margineulima)? sp. (fig. 88) 1933 eulima sp. – ravn, p. 41. material. only the illustrated specimens are known. remarks. ravn (1933, p. 42) stated that the specimens were rather similar to the subgenus margineulima, but the poor material precludes a precise determination. fig. 88. eulima (margineulima?) sp. a: mguh 31050 (ex rkz 45a), height 9.4 mm, width 3.7 mm. b: mguh 31051 (ex rkz 45b), height 10.4 mm, width 4.2 mm. genus melanella bowdich 1822 type species. melanella dufresnii bowdich 1822 melanella sp. (fig. 89) material. only the illustrated specimen is known. remark. the species has a relatively low last whorl, a slightly twisted columella and a rather narrow spout anteriorly. fig. 89. melanella sp. mguh 31052 (ex rkz 47), height 4.7 mm, width 1.5 mm. family aclididae g.o. sars 1878 genus (graphis) jeffreys 1867 type species. graphis unicus (montagu 1803). graphis danica ravn 1933 (fig. 90) 1933 aclis (graphis) danica ravn, p. 38, plate 3, figs 6a, b. type material. holotype mguh 3139. material. the species is rather common. 71 fig. 90. graphis danica ravn 1933. mguh 31053 (ex rkz 51), height 3.7 mm, width 0.9 mm. graphis sp. 1 (fig. 91) additional material. arf, 6 specimens (øsm-1005925043); isl, 1 specimen; mno, 3 specimens. the species is rare. remarks. this species differs from graphis danica in having a distinct undulating suture, caused by the coarser axial ribs. these ribs are more flexuous and on the transition to the base they have an angulation, superficially giving the appearance of a basal disc. rosenkrantz noted that the species might be a scalarid. however, the species has a protoconch very similar to the preceding species, the same slender outline and no real basal disc and we thus prefer to interpret the species as graphis sp. fig. 91. graphis sp. 1. a: mguh 31054 (ex rkz 50a), height 4.3 mm, width 1.1 mm. b: mguh 31055 (ex rkz 50b), height 4.0 mm, width 0.9 mm. a b graphis? sp. 2 (fig. 92) material. only the illustrated specimen is known. remarks. this and the following species resemble the smooth species from the french eocene, as illustrated by gougerot & le renard (1987). unfortunately, the aperture is not preserved. fig. 92. graphis? sp. 2. mguh 31055 (ex rkz 127), height 4.3 mm, width 1.2 mm. 7272 graphis? sp. 3 (fig. 93) material. only the illustrated specimen is known. remarks. the smooth species differs from the previous species by having lower whorls. fig. 93. graphis? sp. 3 mguh 31057 (ex rkz 128), height 5.1 mm, width 1.5 mm. superfamily triphoroidea gray 1847a family triphoridae gray 1847a subfamily triphorinae gray 1847a genus epetrium harris & burrows 1891 type species. triforis grignonensis deshayes 1866. epetrium? cretacea (ravn 1933) (fig. 94) 1933 triforis (epetrium) cretacea ravn, p. 53, plate 5, figs 12a, b, 13a, b. 1998 epetrium? cretacea (ravn 1933) – nützel, p. 127, plate 18, figs k–n. type material. holotype mguh 3174, paratype mguh 3175. additional material. gm, 74 specimens; arf, 116 specimens; isl, 706 specimens; mno, 140 specimens; ebp, 216 specimens. smf 311770, 344078, 150 specimens. rkz 238, 239. the species is abundant. remarks. the species is characterised by having about 17 knobs on each whorl and three visible spirals. of these, the adapical is weak. a fourth spiral is covered by the following whorl. nützel (1998, p. 126) stated that the protoconch of triforis grignonensis deshayes 1866, the type species of epetrium, is unknown and that several species differing from the type species have been assigned to epetrium. he also stated that epetrium? cretacea has only two knob-bearing spirals and e.? faxensis (= triphora (ogivia) faxensis) has a different teleoconch sculpture. based on these observations, e. crassigranulata and e. cretacea were only tentatively referred to the genus epetrium. the study concluded, however, that these two species are the oldest known triphorinae with a preserved protoconch (nützel 1998, p. 126). fig. 94. epetrium? cretacea (ravn 1933). mguh 31058 (ex rkz 238), height 7.8 mm, width 2.1 mm. epetrium? crassigranulata (ravn 1933) (fig. 95) 1933 triforis (epetrium) crassigranulata ravn, p. 54, plate 6, figs 1a, b, 2a, b. 1998 epetrium? crassigranulata ((ravn 1933) – nützel, p. 126, plate 18, fig. j. 73 type material. holotype mfuh 3281, paratype mguh 3282. additional material. gm, 12 specimens; arf, 2 specimens; isl, 5 specimens; smf 311 769. the species is rare. ba fig. 95. epetrium? crassigranulata (ravn 1933). mguh 31059 (ex isl), height 5.0 mm, width 1.7 mm. epetrium? separabilis (ravn 1933) (fig. 96) 1933 triforis (epetrium) separabilis ravn, p. 53, plate 5, figs 17a, b. type material. holotype mguh 3179. additional material. isl, 9 specimens; ebp, one specimen and a fragment; mno, 1 specimen. the species is rare. remarks. this species differs in having four spirals on the teleoconch whorls. the complete specimen rkz 245 consists of five convex protoconch whorls and about eight almost flat and relatively low teleoconch whorls. the whorls have three knob-bearing spirals which are all rather weak. the number of knobs is about 20. the aperture is rounded rectangular. fig. 96. epetrium? separabilis (ravn 1933). a: mguh 31060 (ex rkz 245), height 4.6 mm, width 1.4 mm. b: mguh 31061 (ex rkz 244), height 4.3 mm, width 1.7 mm. epetrium? sp. (fig. 97) additional material. isl, 3 specimens; smf 344080, 1 specimen. the species is very rare. description. the shell is sinistral and slender conical. no specimens with protoconch preserved have been found. the whorls are flat and separated by a distinct suture. the last whorl equals less than 0.1 of the total shell height and the height/width ratio is more than 5. there are three spiral ribs of almost equal intensity and they are a little raised and prominent in lateral view. the two adapical spiral ribs are close-set and separated by a narrow furrow, whereas spirals number two and three in abapical direction are separated by a wider furrow. the flat and smooth base is demarcated by a fourth rather strong and smooth spiral, which is more or less covered by the following whorl. there are c. 18 axial ribs, which are opisthocline. on the three spiral ribs they cause more or less rhomboidal knobs. the aperture is quadrangular with rounded corners and the anterior canal is tubular and closed, slightly turned backwards. the posterior canal has an oblique tube and is situated near the adapical suture. one specimen (fig. 97b) differs by having almost flat whorls and almost equal and equidistant spiral ribs. 7474 remarks. the species differs from epetrium? cretacea by having three spiral ribs of equal strength, of which the adapical two spiral ribs are more close-set, flat whorls and a more slender outline. epetrium? separabilis has four spirals with knobs. the knobs are increasing in strength in abapical direction on the three upper spirals, whereas the spiral above the apical suture has much smaller knobs. fig. 97. epetrium? sp. a: mguh 31062 (ex isl), height 5.2 mm, width 1.7 mm. b: mguh 31063 (ex isl), height 2.8 mm, width 1.1 mm. photos: alexander nützel, münchen. genus triphora blainville 1828 type species. triphora gemmata blainville 1828. subgenus triphora (ogivia) harris & burrows 1891 type species. triphora singularis deshayes 1864. triphora (ogivia) faxensis ravn 1933 (fig. 98) 1933 triphora (ogivia) faxensis ravn, p. 55, plate 5, figs 14a, b. type material. holotype mguh 3176. additional material. arf, 1 specimen (arf 25117); isl, 1 specimen; smf 344079, 1 specimen. the species is very rare. remarks. this species is characterised by having slightly undulating sutures and a canal turned backwards. the whorls have two spiral furrows, of which the adapical is narrow and the abapical is wider. these furrows separate three spiral ribs, of which the adapical one is weak. the middle spiral is the strongest and the abapical is a little weaker. the axial ribs are slightly prosocline and cause knobs on the spiral ribs, especially on the middle spiral. the number of axial ribs is about 14. fig. 98. triphora (ogivia) faxensis ravn 1933. a, b: mguh 3176, height 4.0 mm, 2.3 mm. c: mguh 31064, height 5.5 mm, width 1.4 mm. triphora (ogivia) sp. (fig. 99) 1933 triphora (ogivia) sp. – ravn, p. 55. addtional material. only one specimen was mentioned by ravn (mguh 31075, ex gm 1991.5752); isl, 7 specimens; mno, 6 specimens. the species is rare. remarks. this species has slightly convex whorls and four spirals. the three upper spirals are increasing in strength abapically, and the third in abapical direction is somewhat stronger than the other spirals. the abapical spiral is much weaker. the number of axial ribs is about 18 and they cause knobs on the spirals, especially on the third spiral, where the knobs are large and oblique. 75 fig. 99. triphora (ogivia) sp. a: mguh 31065 (ex rkz 239), height 2.4 mm, width 1.0 mm. b: mguh 31066 (ex gm 1991.5752), height 2.8 mm, width 1.4 mm. a b triphoridae gen. et sp. indet. 1 (fig. 100) additional material. isl, 1 specimen. the species is very rare. remarks. the fragmentary specimen rkz 240 consists of the last three teleoconch whorls, which have three spirals with 16–18 rectangular knobs. of these, the abapical spiral is the strongest. a fourth smooth spiral demarcates the base. the fragmentary specimen rkz 247 has a little more than one whorl preserved. there are three broad spiral bands, which have about 16 weak rectangular knobs. the axial ribs are prosocline. on the base, six further spirals are present. fig. 100. triphoridae gen. et sp. indet. 1. a: mguh 31067 (ex rkz 240), height 3.8 mm, width 1.7 mm. b, c: mguh 31068 (ex rkz 247), height 3.8 mm, width 2.6 mm. triphoridae gen. et sp. indet. 2 (fig. 101) additional material. isl, 9 specimens. the species is rare. remarks. the fragmentary specimens have almost smooth and relatively low whorls. abapically a smooth spiral band is present. the aperture is rather narrow. the fragmentary specimen rkz 242 consists of about seven almost flat and low whorls, which wear three weak spirals. the spirals wear knobs. a large specimen has been found as an external impression in the coral limestone. it has 20 whorls and the sculpture fades away completely on the younger whorls. fig. 101. triphoridae gen. et sp. indet. 2. a: mguh 31069 (ex rkz 242), height 10.3 mm, width 4.0 mm. b: mguh 31070 (ex rkz 243), height 4.8 mm, width 2.2 mm. c: mguh 31071 (ex rkz 246), height 4.2 mm, width 2.3 mm. d: mguh 31072 (ex rkz 241), height 4.3 mm, width 3.2 mm. family cerithiopsidae h. & a. adams 1853 subfamily cerithiopsinae h. & a. adams 1853 genus cerithiopsis forbes & hanley 1851 type species. cerithiopsis tubercularis montagu 1803. 7676 cerithiopsis unisulcata ravn 1933 (fig. 102) 1933 cerithiopsis unisulcata ravn, p. 48, p. 50, plate 5, figs 3a, b, 4a, b. type material. holotype mguh 3166, paratype mguh 3165. additional material. gm, 53 specimens; smf 344068, 109 specimens. the species is very common. fig. 102. cerithiopsis unisulcata ravn 1933. a: mguh 31074 (ex rkz 195), height 7.3 mm, width 2.2 mm. b: mguh 31075 (ex rkz 196a), height 6.9 mm, width 1.8 mm. cerithiopsis aff. unisulcata ravn 1933 (fig. 103) material. the two illustrated specimens are the only specimens known. remarks. the species differs from c. unisulcata by having the strongest knobs on the abapical spiral instead of on the adapical spiral. fig. 103. cerithiopsis aff. unisulcata ravn 1933. a: mguh 31076 (ex rkz 194), height 1.4 mm, width 0.7 mm. b: mguh 31077 (ex rkz 212), height 3.5 mm, width 0.9. mm cerithiopsis rosenkrantzi (ravn 1933) (fig. 104) 1933 tenuiscala rosenkrantzi ravn, p. 39 (partim), plate 3, figs 8a, b (non figs 9a, b, 10a, b = vatopsis metaxiformis). 1998 cerithiopsis cf. rosenkrantzi ravn – nützel, p. 103, plate 16, fig. p. type material. holotype mguh 3142. additional material. gm, 2 specimens; smf 311755, 1 specimen; smf 344074 and 344075, 3 specimens; mno, 1 specimen. the species is rare. remarks. ravn selected figs 8a–b as holotype, but noted that the specimen on figs 9a–b has a more conoid protoconch. in his description he noted that the teleoconch whorls have three rather strong spirals, crossed by 10–11 strong axial ribs. nützel (1998, p. 103) described, based on two specimens, the species and stated that the protoconch has six smooth convex whorls and the teleoconch has two spirals. nützel (1998, p. 111, plate 16, fig. k) established vatopsis metaxiformis, which has a protoconch with spiral ribs and axial ribs. the protoconch whorls are slightly angular at the two abapical spiral ribs. the teleoconch whorls have three strong spiral ribs and c. 10 77 axial ribs. it is obvious that ravn included this species in tenuiscala rosenkrantzi. a study of the material in gm showed that almost all of the 60 specimens, labelled tenuiscala rosenkrantzi, in fact are vatopsis metaxiformis. only the holotype and two specimens are genuine tenuiscala rosenkrantzi. furthermore, the state of conservation of the type material in the type collection of geological museum is very poor. fig. 104. cerithiopsis rosenkrantzi (ravn 1933). smf 311755, height 1.4 mm, width 0.4 mm. photo: alexander nützel, münchen. cerithiopsis bruennichi ravn 1933 (fig. 105) 1933 cerithiopsis brünnichi ravn, p. 49, p. 50, plate 4, figs 14a, b, 15a, b. type material. holotype mguh 3162, paratype mguh 3161. additional material. gm (ravn), 2 specimens; gm 1991.4159, 3 specimens; gm 1991.4160, 9 specimens; gm 1991.4161, 2 specimens; arf, 24 specimens (øsm10059-25031); isl, 4 specimens. the species is rather common. remarks. the species has 20–22 axial ribs, whereas zaclys? selandica has only c. 15 axial ribs. furthermore, the protoconchs are different. fig. 105. cerithiopsis bruennichi ravn 1933. a: mguh 3162, height 5.7 mm, width 1.9 mm. b: mguh 3161, height 0.9 mm, width 0.4 mm. cerithiopsis sp. 1 (fig. 106) material. only the illustrated specimen is known (ex isl). remarks. the protoconch has opisthocline axial ribs and a teleoconch with three primary spirals, of which the abapical is weak and situated immediately above the suture. between the two adapical spirals a secondary spiral is inserted. the spirals run across about 20 slightly opisthocline axial ribs. the aperture is rounded rectangular with a short canal, which is turned to the left. 7878 fig. 106. cerithiopsis sp. 1. mguh 31078 (ex isl), height 2.8 mm, width 1.0 mm. cerithiopsis sp. 2 (fig. 107) material. only the illustrated specimen is known. remarks. this species has two spirals, crossed by 14 axial ribs. at the intersections, large oblong knobs occur. zaclys? selandica (nützel 1998, p. 104, plate 16, fig. o) has considerably weaker knobs. cerithiopsis sp. 13 has two spirals, crossed by 16 axial ribs, but differs by having a smooth spiral band under the two knob-bearing spirals. fig. 107. cerithiopsis sp. 2. mguh 31079 (ex rkz 72), height 3.2 mm, width 0.9 mm. cerithiopsis sp. 3 (fig. 108) material. only the illustrated specimen is known. remarks. the species has almost flat whorls, bearing three broad spirals of almost the same strength. the axial ribs (about 14 on each whorl) are weaker and divide the spirals into oblong knobs. the protoconch whorls are smooth, but only the last four whorls are preserved. fig. 108. cerithiopsis sp. 3. mguh 31080 (ex rkz 190), height 5.7 mm, width 1.4 mm. cerithiopsis sp. 4 (fig. 109) additional material. gm 1991.4184, 2 specimens; gm 1991.4197, 2 specimens; gm 1991.4173, 1 specimen; gm 1991.4184, 1 specimen; smf 344070, 1 specimen. the species is rare. remarks. on the first teleoconch whorls there are three spiral bands, each bearing 18–24 knobs. the knobs on the adapical spiral are higher than the other knobs. on a few specimens, the spirals on the last teleoconch whorls are divided into six spirals, probably due to growth errors. 79 fig. 109. cerithiopsis sp. 4. a: mguh 31081 (ex rkz 201a), height 6.7 mm, width 1.8 mm. b: mguh 31082 (ex rkz 201c), height 4.9 mm, width 1.3 mm. c: mguh 31083 (ex rkz 201b), height 10.2 mm, width 4.2 mm. d: mguh 31084 (ex rkz 200), height 4.7 mm, width 1.6 mm. cerithiopsis sp. 5 (fig. 110) additional material. gm 1991.4185, 3 specimens. the species is very rare. remarks. there are three flat spiral bands on the teleoconch, and the adapical is considerably weaker. there are 12–13 flat axial ribs, separated by narrower interspaces. fig. 110. cerithiopsis sp. 5. mguh 31085 (ex rkz 204), height 4.0 mm, width 0.9 mm. cerithiopsis sp. 6 (fig. 111) additional material. gm 1991.4184, 1 specimen. the species is very rare. remarks. the teleoconch has two spirals and c. 14 axial ribs. the smooth, multispiral protoconch has slightly angular whorls. c. rosenkrantzi has a rather similar teleoconch, but a smooth protoconch with regularly convex whorls. fig. 111. cerithiopsis sp. 6. mguh 31086 (ex rkz 208), height 3.2 mm, width 0.8 mm. cerithiopsis sp. 7 (fig. 112) additional material. gm 1991.4194, 1 specimen. the species is very rare. remarks. the multispiral protoconch has opisthocline axial ribs on the terminal two whorls and no spirals. the teleoconch has slightly convex whorls and three spirals, bearing c. 16 knobs on each whorl. 8080 fig. 112. cerithiopsis sp. 7. mguh 31087 (ex rkz 210), height 3.9 mm, width 1.1 mm. cerithiopsis sp. 8 (fig. 113) additional material. gm 1991.4185, 3 specimens. gm 1991.4201, 3 specimens. the species is rare. remarks. the teleoconch has almost flat whorls, bearing about 16 rounded knobs. there is a subsutural smooth band. the base is flat. fig. 113. cerithiopsis sp. 8. mguh 31088 (ex rkz 211), height 6.7 mm, width 1.2 mm. cerithiopsis sp. 9 (fig. 114) additional material. gm 1991.4191, 1 specimen. the species is very rare. remarks. the teleoconch has almost flat whorls, bearing 16 coarse knobs. there are three spirals on the last whorls, of which the adapical is much weaker than the two other spirals. the shell is less slender than other species of cerithiopsis. fig. 114. cerithiopsis sp. 9. mguh 31089 (ex rkz 213), height 3.2 mm, width 1.2 mm. cerithiopsis sp. 10 (fig. 115) additional material. gm 1991.4186, 1 specimen. the species is very rare. remarks. the teleoconch has convex whorls, bearing c. 16 coarse knobs. abapically there is a smooth sutural band. 81 fig. 115. cerithiopsis sp. 10. mguh 31090 (ex rkz 214), height 2.9 mm, width 0.7 mm. genus zaclys finlay 1926 type species. cerithiopsis sarissa murdoch 1905. zaclys? selandica (ravn 1933) (fig. 116) 1933 cerithiopsis selandica ravn, p. 51, plate 5, figs 10a, b (partim, non figs 9a, b = zaclys? nuetzeli n. sp.). 1998 zaclys? sp. – nützel, p. 103, plate 16, fig. o. type material. holotype mguh 3171. additional material. gm 1991.4159, 3 specimens; gm 1991.4160, 2 specimens; smf 311753, 2 specimens; smf 344075, 2 specimens. the species is rare. discussion: ravn (1933) illustrated two specimens, of which the specimen illustrated on plate 5, figs 9a–b (mguh 3171) was designated as holotype. in the rosenkrantz files of drawings, the type material is also illustrated. the specimen mguh 3171 has a protoconch, consisting of c. five whorls, of which the terminal has two very fine spiral ribs, causing a slight angulation of the whorl. adapically fine axial ribs are suggested, but only visible near the adapical suture. the teleoconch of mguh 3171 has c. 15 axial ribs. mguh 3172 has a considerably larger protoconch with axial ribs suggested and no spiral ribs on the last whorl. the teleoconch has 12 axial ribs on each whorl. thus, ravn’s description does not match the holotype and it is obvious that ravn’s material contains two species. of the two illustrated types mguh 3172 matches the description in ravn (1933), but according to iczn (1999, article 70.3), mguh 3171 is the holotype. for this reason mguh 3172 is established as the new species zaclys? nuetzeli. remarks. nützel (1998, p. 103–104, plate 16, figs m, n, o) recognised two zaclys? species in his material and interpreted cerithiopsis selandica in accordance with ravn’s description. the species is rather similar to cerithiopsis bruennichi (for comparison see remarks to this species). nützel (1998, p. 33 and p. 103) discussed the genus zaclys finlay 1926 from the danian of faxe and concluded that the morphology of the two danish species were very similar to this genus. however, as the characters of the radula are unknown, he assigned the danish species to zaclys with a query. fig. 116. zaclys? selandica (ravn 1933). mguh 3171, height 2.2 mm, width 0.8 mm. zaclys? nuetzeli n. sp. (fig. 117) 1933 cerithiopsis selandica ravn, p. 51 (partim, non figs 9a, b). 1998 zaclys? selandica ravn – nützel, p. 104, plate 16, figs m, n (non ravn). type locality. faxe quarry, sjælland, denmark. type stratum. faxe formation, middle danian, paleocene. 8282 type material. holotype mguh 3172. additional material. smf 311 754; rkz 202; arf, 2 specimens (øsm-10059-25031 and 25032); isl, 13 specimens.; gm 1991.4158, 1 specimen; gm 1991.4159, 1 specimen; gm 1991.4160, 10 specimens. the species is rather common. derivation of name. the species is named in honour of alexander nützel, who first recognised the two zaclys? species in the faxe material. diagnosis. a zaclys? with a protoconch consisting of c. five whorls with axial ribs and no spirals on the last whorl. the teleoconch has c. 12 axial ribs. description. the shell is slender and turriculate with a height/width ratio of c. 3.0. the protoconch has six convex whorls, separated by a distinct suture. the first three whorls are smooth and the two terminal whorls have a spiral keel under the middle of the whorl and axial ribs at the adapical suture. the transition into the teleoconch is sharp. the teleoconch whorls are convex and separated by a distinct suture. the spiral ornament consists of four spirals, of which the adapical is weak and situated under the adapical suture. the abapical spiral is weak and situated above the abapical suture and partly covered by the following whorl. the two medium spirals are of almost equal strength and run across 12 axial ribs on each whorl. the knobs on the two spirals are rather weak. the base is smooth. the aperture is rounded rectangular and the canal is short. remarks. the differences to zaclys? selandica have been discussed above. fig. 117. zaclys? nuetzeli n. sp. a: mguh 31091 (ex rkz 205), height 3.1 mm, width 0.9 mm. b: mguh 31092 (ex rkz 202), height 3.9 mm, width 1.0 mm. c: mguh 3172, height 2.5 mm, width 0.8 mm. d: mguh 31093 (ex rkz 209), height 3.3 mm, width 0.9 mm. genus eocolina chavan 1952 type species. cerithium munieri deshayes 1864. eocolina sp. (fig. 118) material. only the two illustrated specimens are known. the species is very rare. remarks. the species differs from the preceding species by having a more slender protoconch instead of a broad conical protoconch, which consists of fewer whorls. the protoconch whorls have very weak spiral riblets. furthermore, the species has about 30 axial ribs, which are sharper and slightly wavy. the spirals and axial ribs together result in a rectangular cancellation. 83 fig. 118. eocolina sp. a: mguh 31094 (ex rkz 197), height 4.2 mm, width 1.3 mm. b: mguh 31095 (ex rkz 215), height 4.6 mm, width 4.0 mm. genus retilaskeya marshall 1978 type species. retilaskeya zelandica marshall 1978. retilaskeya ravni nützel 1998 (fig. 119) 1998 retilaskeya ravni nützel, p. 114, plate 16, figs h, i. type material. smf 311 748. additional material. gm 1991.4090, 5 specimens; gm 1991.4121, 1 specimen; gm 1991.4146, 4 specimens and 1 dubious; gm 1991.4158, 5 specimens; gm 1991.4160, 1 specimen; gm 1991.4191, 1 specimen; arf, 1 specimen (øsm-10059-25039); isl, 5 specimens; ebp, 2 specimens. the species is rather common. fig. 119. retilaskeya ravni nützel 1998. holotype smf 311 748, height 2.4 mm, width 0.7 mm. photo: alexander nützel, münchen. retilaskeya sp. 1 (fig. 120) additional material. gm 1991.4127, 1 specimen. gm 1991.4163, 2 fragments. the species is very rare. remarks. the protoconch is highly conical with opisthocline axial ribs and is similar to the protoconch of retilaskeya. the teleoconch has two strong spirals with c. 14 strong knobs. there is a depression with two furrows between the spirals. fig. 120. retilaskeya sp. 1. mguh 31096 (ex gm 1991.4127), height 3.0 mm, width 0.8 mm. 8484 retilaskeya sp. 2 (fig. 121) additional material. gm 1991.4162, 2 juvenile specimens, 9 fragments. the species is rare. remarks. the protoconch has five whorls with opisthocline axial ribs, which is similar to the protoconch of retilaskeya. the teleoconch has two spirals with c. 16 knobs, separated by a slight depression. the knobs on the adapical spiral are oblong and orthocline, whereas the abapical knobs are opisthocline. fig. 121. retilaskeya sp. 2. mguh 31097 (ex gm 1991.4162), height 3.1 mm, width 0.9 mm. genus krachia baluk 1975 type species. cerithiopsis (krachia) korytnicensis baluk 1975. krachia sp. (fig. 122) material. the species is rather common. fig. 122. krachia sp. mguh 31098. (ex rkz 206), height 13.7 mm, width 4.3 mm. genus specula finlay 1926 type species. cerithiopsis styliformis suter 1908. specula angustisulcata (ravn 1933) (fig. 123) 1933 newtoniella angustisulcata ravn, p. 46, plate 5, figs 1a, b, 2a, b. 1998 specula? angustisulcata (ravn 1933) – nützel, p. 118, plate 16, figs t–w. type material. holotype mguh 3163. additional material. gm, 158 specimens; smf 311 759, 311 760. the species is very common. fig. 123. specula angustisulcata (ravn 1933). mguh 31099 (ex rkz 198), height 4.3 mm, width 1.1 mm. 85 subfamily aliptinae marshall 1978 genus cerithiopsidella bartsch 1911 type species. cerithiopsis cosmia bartsch 1907. cerithiopsidella trinodosa (ravn 1933) (fig. 124) 1933 cerithiopsis trinodosa ravn, p. 49, p. 50, plate 4, figs 5a, b. 1998 cerithiopsidella trinodosa (ravn 1933) – nützel, p. 109, plate 16, fig. g. type material. holotype mguh 3152. additional material. smf 311 764; gm, 3 specimens (1991.4191 and 1991.4194); arf, 2 specimens (øsm10059-25024 and 25090); ebp, 1 specimen. the species is rare. fig. 124. cerithiopsidella trinodosa (ravn 1933). a: mguh 31100 (ex rkz 186), height 4.8 mm, width 1.3 mm. b: mguh 3152, height 5.8 mm, width 2.1 mm. cerithiopsidella sp. (fig. 125) additional material. gm 1991.4150, 2 specimens. the species is very rare. remarks. the species is very slender and the teleoconch whorls have two spirals with c. 14 knobs on each whorl. the knobs on the adapical spiral are considerably stronger than the knobs on the lower spiral. under the adapical suture there is an almost flat band without sculpture. fig. 125. cerithiopsidella sp. øsm-10059-25033 (ex arf), height 6.8 mm, width 1.1 mm genus vatopsis gründel 1980 type species. cerithium bimonilifera sandberger 1858. vatopsis metaxiformis nützel 1998 (fig. 126) 1933 tenuiscala rosenkrantzi ravn, p. 39 (partim), plate 3, figs 9a, b (non figs 8a, b, 10a, b = cerithiopsis rosenkrantzi) 1998 vatopsis metaxiformis nützel, p. 111, plate 16, fig. k. type material. smf 311 750, 6 specimens. additional material. smf 344073, 115 specimens; mguh 3143, rkz 203. 8686 remarks. the species is very common. fig. 126. vatopsis metaxiformis nützel 1998. a: mguh 31102 (ex rkz 62), height 2.1 mm, width 0.6 mm. b: mguh 31103 (ex rkz 203), height 4.8 mm, width 0.8 mm. subfamily seilinae golikov & starobogatov 1975 genus variseila dockery 1993 type species. cerithiopsis meeki wade 1926. variseila eocostata nützel 1998 (fig. 127) 1998 variseila eocostata nützel, p. 105, plate 16, fig. e. type material. smf 311 744. additional material. rkz 199. gm, 6 specimens (1991.4129 1991.4153 and 1991.4160). the species is rare. remarks. the teleoconch of this species is rather similar to thereitis tricingulata, but the protoconchs of the two species are different. fig. 127. variseila eocostata nützel 1998. mguh 31104 (ex rkz 199), height 5.1 mm, width 1.2 mm. variseila fissicosta (ravn 1933) (fig. 128) 1933 newtoniella fissicosta ravn, p. 48, plate 4, figs 6a, b, 10a, b. type material. holotype mguh 3153, paratype mguh 3157. additional material. rkz 69; gm, 2 specimens (1991.4116); arf, 66 specimens (øsm-10059-25040 and øsm-10059-25041); isl, 5 specimens. the species is common. remarks. this species has a characteristic sculpture, consisting of three spirals, of which the abapical is much stronger than the two other spirals and subdivided into three fine spirals. 87 fig. 128. variseila fissicosta (ravn 1933). a: mguh 31105 (ex rkz 69), height 5.5 mm, width 2.7 mm. b: mguh 31106 (ex rkz 189), height 3.0 mm, width 0.8 mm. c: mguh 3157, height 2.3 mm, width 0.9 mm. d: mguh 3153, height 2.6 mm, width 2.3 mm. variseila sp. 1 (fig. 129) additional material. gm, 1 specimen (1991.4099). the species is very rare. remarks. this species has slightly convex whorls and five flat spiral bands. fig. 129. variseila sp. 1. mguh 31107 (ex rkz 191), height 7.3 mm, width 1.6 mm. variseila sp. 2 (fig. 130) material. one complete specimen and one fragment are known. the species is very rare. description. the shell has a height of 3.0 mm and provides about four protoconch whorls and nine teleoconch whorls. the first two protoconch whorls are quickly increasing in diameter, whereas the terminal two are of almost the same diameter. the teleoconch has four spirals of almost equal strength. the aperture is narrow and subcircular, with a short canal. fig. 130. variseila sp. 2. mguh 31108 (ex gm 1991.4111), height 3.0 mm, width 0.6 mm. 8888 genus seila a. adams 1861 type species. triphoris dextroversus a. adams & reeve 1850. seila sp. (fig. 131) material. only the illustrated specimen is known. remarks. the multispiral protoconch consists of six convex whorls. the teleoconch has three spirals, of which the adapical two are of the same strength and the abapical one is the strongest. fig. 131. seila sp. mguh 31109 (ex gm 1991.4096), height 2.7 mm, width 0.8 mm. subgenus seila (notoseila) finlay 1926 type species. cerithium terebelloides hutton 1873. seila (notoseila) sp. 1 (fig. 132) additional material. gm 20 specimens (gm 1991.4122 1991.4123 1991.4124 and 1991.4134); arf, 1 specimen (øsm-10059-25028). the species is rather common. remarks. this species has completely flat whorls and four flat spiral bands. fig. 132. seila (notoseila) sp. 1. mguh 31110 (ex rkz 184), height 5.3 mm, width 0.8 mm. seila (notoseila) sp. 2 (fig. 133) material. only the illustrated specimen is known. remarks. this species has a protoconch, consisting of about six convex and smooth whorls. the spiral ornamentation of the teleoconch consists of four spirals, of which the two adapical are weaker than the two abapical spirals. fig. 133. seila (notoseila) sp. 2. mguh 31111 (ex rkz 188), height 3.6 mm, width 0.8 mm. 89 genus thereitis le renard 1997 type species. seila (notoseila?) angusta tembrock 1964. thereitis tricingulata (ravn 1933) (fig. 134) 1933 cerithiopsis tricingulata ravn, p. 52, plate 5, figs 11a, b. 1998 tembrockia tricingulata (ravn 1933) – nützel, p. 105, plate 16, fig. e. type material. holotype mguh 3173. additional material. gm, 34 specimens (gm 1991.4089 1991.4090 1991. 4134 1991.4135 1991.4151 1941.4152 and 1991.4153); arf, 2 specimens (øsm-10059-25040 and 25821); isl, 1 specimen; smf 311 745, 1 specimen. the species is rather common. fig. 134. thereitis tricingulata (ravn 1933). mguh 3173, height 5.3 mm, width 0.8 mm. family newtoniellidae korobkov 1955 subfamily newtoniellinae korobkov 1955 genus cerithiella verrill 1882 type species. cerithium metula lovén 1846. cerithiella faxensis (ravn 1933) (fig. 135) 1933 newtoniella faxensis ravn, p. 48, plate 4, figs 3a, b, 9a, b. type material. holotype mguh 3156, paratype mguh 3156. additional material. gm (ravn), 3 specimens; arf, 77 specimens (øsm-10059-25062, 25069 and 25099); isl, 5 specimens. the species is common. fig. 135. cerithiella faxensis (ravn 1933). mguh 31113 (ex isl), height 8.8 mm, width 3.1 mm. cerithiella fenestrata (ravn 1902b) (fig. 136) 1902 cerithium fenestratum ravn, p. 222, plate 1, figs 20, 21. 1933 “cerithium” fenestratum ravn, – ravn, p. 45. type material. syntypes mguh 94 and mguh 95. material. specimens of this large species are rather common. remarks. ravn noted that the sculpture was rather similar to the sculpture of the genus newtoniella (= cerithiella), but the large size made an assignment to this genus questionable. however, the general outline and the aperture with the twisted columella are comparable to cerithiella. 9090 a b fig. 136. cerithiella fenestrata (ravn 1902b). mguh 31114, height 13.8 mm, width 5.3 mm. cerithiella sp. (fig. 137) additional material. gm 1991.4149, 12 specimens; smf 344069, 5 specimens; smf 344071, 1 specimen; isl, 1 specimen. the species is rare. remarks. the species has two spirals, separated by a depression. the spirals have about 20 rounded knobs, but the knobs on the adapical spiral are larger and indistinct. fig. 137. cerithiella sp. a: mguh 31115 (ex rkz 196c), height 4.9 mm, width 1.4 mm. b: mguh 31116 (ex rkz 196b), height 5.2 mm, width 2.0 mm. genus trituba jousseaume 1884 type species. triforis bitubulatus baudon 1856. trituba obliquecostulata (ravn 1933) (fig. 138) 1933 cerithiopsis obliquecostulata ravn, p. 52, plate 4, figs 12a, b, 13a, b. type material. holotype mguh 3159, paratype mguh 3160. additional material. gm, 6 specimens; arf, 36 specimens (øsm-10059-25034 and 25035); isl, 22 specimens; mno, 4 specimens; ebp, 18 specimens. the species is common. remarks. the name triforis deshayes 1834 is now generally considered as a junior synonym of triphora blainville 1828, as deshayes misspelled triphora (gofas & rosenberg 2014). gougerot & le renard (1980) and marshall (1980) used the name triforis for this dextral genus. this dextral species matches with regard to protoconch and teleoconch sculpture and aperture the genus trituba very well. the specimen øsm-1005991 25044 (fig. 138c) has a well-preserved aperture, which definitely confirms the assignment to the genus trituba. fig. 138. trituba obliquecostulata (ravn 1933). a: mguh 3159, height 2.0 mm, width 0.7 mm. b: mguh 3160, height 2.7 mm, width 1.3 mm. c: øsm-10059-25044 (ex arf), height 5.2 mm, width 1.8 mm. subfamily eumetulinae golikov & starobogatov 1975 genus eumetula thiele 1912 type species. eumetula dilecta thiele 1912. eumetula multituberculata nützel 1998 (fig. 139) 1998 eumetula? multituberculata nützel, p. 117, plate 16, figs q, r. type material. smf 311 756. additional material. smf 344076, 1 specimen; gm 1991.4102, 2 specimens; gm 1991.4107, 3 specimens; gm 1991.4108, 3 specimens; gm 1991.4109, 3 specimens; gm 1991.4145, 1 specimen; gm 1991.4200, 2 specimens; isl, 5 specimens; mno, 1 specimen. the species is rather common. remarks. the species differs from eumetula jenseni by the protoconch, which has axial ribs. fig. 139. eumetula multituberculata nützel 1998. mguh 31117 (ex gm 1991.4145), height 2.5 mm, width 0.7 mm. eumetula jenseni (ravn 1933) (fig. 140) 1933 cerithiopsis jenseni ravn, p. 49, p. 51, plate 4, figs 11a, b. 1998 eumetula? jenseni (ravn 1933) – nützel, p. 118, plate 16, fig. s. type material. holotype mguh 3158. additional material. gm (ravn 1933), 6 specimens; gm 1991.4200, 1 specimen; 1991.4143, 1 specimen; gm 1991.4145, 6 specimens; gm 1991.4146, 48 specimens; gm 1991.4147, 9 specimens; gm 1991.4165, 2 specimens; arf, 18 specimens (øsm-10059-25025 and 25055); isl, 20 specimens; mno, 10 specimens; smf 311 758, 2 specimens. the species is common. 9292 fig. 140. eumetula jenseni (ravn 1933). mguh 3158, height 5.8 mm, width 2.2 mm. eumetula sp. 1 (fig. 141) additional material. gm 1991.4182, 3 specimens; gm 1991.4180, 10 fragments. the species is rare. remarks. the teleoconch has two spirals with about 12 rounded knobs on each whorl. of these, the abapical is much stronger and protruding than the adapical one, which is situated on a straight part of the whorl. the protoconch is preserved on two specimens and is similar to the protoconch on the genus eumetula, which has axial ribs. fig. 141. eumetula sp. 1. mguh 31118 (ex gm 1991.4182), height 2.1 mm, width 0.7 mm. eumetula sp. 2 (fig. 142) material. only the illustrated specimen is known. remarks. the species has three spirals with knobs, of which spiral number two is considerably stronger than the two other spirals. the number of knobs is c. 12 on each whorl. fig. 142. eumetula sp. 2. mguh 31119 (ex gm 1991.4195), height 3.9 mm, width 1.2 mm. eumetula sp. 3 (fig. 143) material. only the illustrated specimen is known. remarks. the specimen lacks the first protoconch whorls and is characterised by having four spirals and about 25 axial ribs, which cause rectangular knobs on the spirals. the adapical spiral is weaker than the other spirals. on the last whorl the axial ribs completely disappear. the species eumetula jenseni and e. multituberculata have a rather similar teleoconch sculpture, but have a different aperture with a straight columella. as the protoconch is not completely preserved and the columella is concave, the assignment to the genus eumetula is uncertain. 93 fig. 143. eumetula sp. 3. mguh 31120 (ex rkz 183), height 5.6 mm, width 1.9 mm. clade neogastropoda cox 1960 superfamily buccinoidea rafinesque 1815 family fasciolariidae gray 1853 subfamily fusininae wrigley 1927 genus fusinus rafinesque 1815 type species. murex colus linnaeus 1758. fusinus sp. 1 (fig. 144) additional material. arf, 5 specimens (øsm-1005925144); isl, 3 specimens. the species is rare. remarks. the species has a spiral ornament consisting of 10–11 flat spiral bands, separated by narrow spiral furrows. fig. 144. fusinus sp. 1. a, b: mguh 31121 (ex rkz 155), height 4.6 mm, width 2.4 mm. c: mguh 31122 (ex rkz 140), height 3.2 mm, width 1.8 mm. fusinus sp. 2 (fig. 145) material. only the illustrated specimen is known. remarks. the species has only three primary spirals, which are situated on the abapical half of the whorl and separated by furrows of almost the same width. the protoconch whorls are more convex than on fusinus sp. 1. 9494 fig. 145. fusinus sp. 2. øsm-10059-25073 (ex arf), height 5.1 mm, width 1.8 mm. genus dolicholatirus bellardi 1886 type species. drilluta communis wade 1916. dolicholatirus sp. 1 (fig. 146) additional material. rkz 170; arf, 1 specimen (øsm10059- 25105); isl, 8 specimens; ebp, 1 specimen. the species is rare. remarks. the shell is slender and has a spiral ornament consisting of seven spiral ribs, separated by wider interspaces. the adapical four spiral ribs are weaker and more close-set than the abapical three spirals. fig. 146. dolicholatirus sp. 1. a: mguh 31124 (ex rkz 170b), height 8.5 mm, width 3.0 mm. b: mguh 31125 (ex rkz 170a), height 8.0 mm, width 3.2 mm. c: mguh 31126 (ex rkz 170d), height 6.5 mm, width 3.0 mm. dolicholatirus sp. 2 (fig. 147) additional material. arf, 1 specimen (øsm-1005925114). the species is very rare. remarks. the shell is less slender than dolicholathyrus sp. 1 and has a spiral ornament consisting of eight spiral bands, separated by narrow furrows. the axial ribs are strong and the distinct growth lines cause flat knobs on the spiral bands. the specimen has apparently no columellar folds, which may be due to the juvenile stage of it. the juvenile specimen of the proceeding species shows no axial ribs, while a larger specimen has the two columellar folds, which characterise the genus. 95 fig. 147. dolicholatirus sp. 2. a, b: mguh 31127 (ex rkz 171c). a: height 4.2 mm, width 1.8 mm. b: height 1.4 mm, width 1.1 mm. c: mguh 31128 (ex rkz 171e), height 3.0 mm, width 1.5 mm. genus conradconfusus snyder 2002 type species. conradconfusus parilis (conrad 1832) conradconfusus parvus (ravn 1933) (fig. 148) 1933 buccinofusus? parvus ravn, p. 63, plate 7, figs 1a, b. 2002 conradconfusus parvus (ravn 1933) – snyder, p. 245. type material. holotype mguh 3193. material. the species is very common, especially as juvenile specimens. remarks. snyder (2002) introduced the new genus taxon conradconfusus, and subsequently assigned the genus to the subfamily fusininae wrigley 1927 of the fasciolariidae gray 1853 (snyder 2003). fig. 148. conradconfusus parvus (ravn 1933). mguh 31129 (ex rkz 250), height 6.3 mm, width 3.0 mm. conradconfusus subglaber (ravn 1933) (fig. 149) 1933 buccinofusus? subglaber ravn, p. 63, plate 7, figs 4a, b. 2002 conradconfusus subglaber (ravn 1933) – snyder, p. 245. type material. holotype mguh 3196. material. the species is very common, especially as juvenile specimens. 9696 fig. 149. conradconfusus subglaber (ravn 1933). mguh 3196, height 7.4 mm, width 4.0 mm. conradconfusus sp. (fig. 150) material. the species is rather common. remarks. the species differs by having less convex whorls than conradconfusus parvus, weaker spirals and axial ribs fading out on the younger whorls. fig. 150. conradconfusus sp. mguh 31130 (ex isl), height 6.3 mm, width 3.0 mm. superfamily muricoidea rafinesque 1815 family muricidae rafinesque 1815 genus pterynotus swainson 1833 type species. murex pinnatus swainson 1822. subgenus pterynotus (pterochelus) jousseaume 1880 type species. murex acanthopterus lamarck 1816. pterynotus (pterochelus) sp. (fig. 151) additional material. gm (rkz 25, rkz 26, rkz 27); arf, 11 specimens (øsm-10059-25077 and 25078); isl, 1 fragment. the species is rare. remarks. rosenkrantz suggested the species name danicus for this species on a drawing, but never published the name. fig. 151. pterynotus (pterochelus) sp. øsm-10059-25078 (ex arf), height 4.5 mm, width 1.9 mm. family costellariidae macdonald 1860 genus vexillum röding 1798 type species. vexillum (vexillum) plicarium linnaeus 1758. vexillum? sp. (fig. 152) 1933 turricula sp., ravn, p. 67, plate 7, figs 9a, b. additional material. mguh 3201; isl, 1 specimen; ebp, 2 specimens; mno, 3 specimens. the species is rare. 97 fig. 152. vexillum? sp. mguh 3201, height 3.0 mm, width 2.0 mm. family mitridae swainson 1829 subfamily mitrininae swainson 1829 genus mitra lamarck 1798 type species. mitra mitra linnaeus 1758. mitra subglabra (ravn 1933) (fig. 153) 1933 turricula (fusimitra) subglabra ravn, p. 65, plate 6, figs 5a, b. type material. mguh 3188. additional material. gm, 13 specimens; arf, 95 specimens (øsm-10059-21426); isl, 12 specimens; mno. the species is very common. fig. 153. mitra subglabra (ravn 1933). mguh 3188, height 8.4 mm, width 3.3 mm. mitra glabra (ravn 1933) (fig. 154) 1933 turricula (fusimitra) glabra ravn, p. 65, plate 7, figs 8a, b. type material. mguh 3197. additional material. gm, 8 specimens; arf, 3 specimens; isl, 3 specimens; mno. the species is rare. fig. 154. mitra glabra (ravn 1933). mguh 3197, height 10.3 mm, width 3.7 mm. 9898 mitra faxensis (ravn 1933) (fig. 155) 1933 turricula (fusimitra) faxensis ravn, p. 66, plate 7, figs 6a, b. type material. mguh 3198. additional material. gm, 5 specimens; arf, 2 specimens (øsm-10059-25147); isl, 1 specimen. the species is rare. fig. 155. mitra faxensis (ravn 1933). mguh 3198, height 8.9 mm, width 3.8 mm. family volutidae rafinesque 1815 subfamily scaphellinae gray 1857 genus scaphella swainson 1832 type species. voluta junonia lamarck 1804. scaphella faxensis (ravn 1902b) (fig. 156) 1902 voluta faxensis ravn, p. 233, plate 3, figs 1–3. 1919 voluta faxensis ravn – nielsen, p. 31. 1933 scaphella faxensis (ravn) – ravn, p. 68. remarks. the species was mentioned by nielsen (1919), but the original material has not been found in the collections of the geological museum, copenhagen. in mno a single internal mould of a juvenile specimen has been found. in general outline, it matches the species well. the species is very rare. fig. 156. scaphella faxensis (ravn 1902b). mguh 31132 (ex mno), height 7.5 mm, width 3.9 mm. family volutomitridae gray 1854 genus conomitra conrad 1865 type species. mitra fusioides lea 1833. conomitra sp. (fig. 157) 1933 conomitra sp. ravn, p. 64, plate 7, figs 3a, b. material. mguh 3195; gm, 2 specimens; arf, 4 specimens (øsm-10059-25086, 25093, 25097 and 25146); isl, 1 specimen; mno, 4 specimens. the species is rare. fig. 157. conomitra sp. mguh 3195, height 11.5 mm, width 4.8 mm. superfamily conoidea fleming 1822 family clavatulidae gray 1853 http://en.wikipedia.org/wiki/conoidea http://en.wikipedia.org/wiki/clavatulidae 99 genus turricula schumacher 1817 type species. turricula flammea schumacher 1817. turricula faxensis (ravn 1902b) (fig. 158) 1902 pleurotoma faxensis ravn, p. 235, plate 3, figs 8, 10. 1933 surcula faxensis ravn, p. 69, plate 7, figs 8a, b. type material. syntypes mguh 127 and mguh 129; mguh 3200. additional material. arf, 11 specimens (øsm-1005925101); isl, 1 specimen; ebp, 1 specimen. the species is rare. remarks. the species is very rare in the baunekule facies. in the coral limestone from other parts of the faxe formation specimens up to 90 mm are found, but the species is rare. fig. 158. turricula faxensis (ravn 1902b). øsm-10059-21049 (ex arf), height 9.5 mm, width 3.3mm. turricula pusilla (ravn 1933) (fig. 159) 1933 surcula? pusilla ravn, p. 70, plate 7, figs 12a, b. type material. holotype mguh 3204. additional material. gm, 3 specimens; arf, 2 specimens; isl, 2 specimens; ebp, 5 specimens; mno, 1 specimen. the species is rare. remarks. on the legend to plate 7, figs 12a, b, ravn misspelled the species name as pussilla. fig. 159. turricula pusilla (ravn 1933). mguh 31135 (ex isl), height 4.0 mm, width 2.1 mm. superfamily cancellarioidea forbes & hanley 1851 family cancellariidae forbes & hanley 1851 subfamily plesiotritoninae beu & maxwell 1987 genus plesiotriton fischer 1884 type species. cancellaria volutella lamarck 1803. plesiotriton steni schnetler & petit 2006 (fig. 160) 2006 plesiotriton steni schnetler & petit, p. 99, plate 1, fig. 3; plate 2, fig. 1. type material. holotype mguh 27344, paratype mguh 27345. additional material. arf, 1 juvenile specimen (øsm10059-25086); isl, 1 fragment; ecs, 1 juvenile specimen (øsm 10061-f 2-46). the species is very rare. 100100 fig. 160. plesiotriton steni schnetler & petit 2006. mguh 27344, height 15.0 mm, width 5.0 mm. subfamily cancellariinae forbes & hanley 1851 genus unitas palmer 1947 type species. cancellaria costulata lamarck 1803. unitas anderseni schnetler & petit 2006 (fig. 161) 2006 unitas anderseni schnetler & petit, plate 1, figs 1a, b; plate 2, figs 2, 7. type material. holotype mguh 27346, paratypes mguh 27347 and mguh 27348. additional material. isl, 2 juvenile specimens; arf, 7 juvenile specimens (øsm-10059-25852); ecs, 1 juvenile specimen (øsm 10061f 2-91); mno, 1 specimen. the species is rare. fig. 161. unitas anderseni schnetler & petit 2006. mguh 27346, height 7.5 mm, width 3.9 mm. unitas aliceae schnetler & petit 2006 (fig. 162) 1933 admete? biplicata (ravn) – ravn, p. 68, plate 6, figs 12a, b. 2006 unitas aliceae schnetler & petit, p. 103, figs 3a, b, 13, 14. type material. holotype mguh 27349, paratypes mguh 27350; mguh 3121, illustrated by ravn (1933, plate 6, figs 12a–b) as admete ? biplicata (ravn). additional material. 2 juvenile specimens, mentioned by ravn (1933, p. 68); arf, 5 specimens (øsm-1005925075); mno, 3 specimens; ebp, 1 specimen. the species is rare. fig. 162. unitas aliceae schnetler & petit 2006. mguh 3121, height 2.6 mm, width 1.8 mm. 101 unitas sp. 1 (fig. 163) 1933 admete (bonellitia) sp. – ravn, p. 68, plate 6, figs 12a, b. 2006 unitas sp. – schnetler & petit, p. 104, figs 15a, b. material. only the illustrated specimen is known. fig. 163. unitas sp. 1. mguh 3192, height 3.2 mm, width 2.1 mm. unitas sp. 2 (fig. 164) additional material. two specimens, gm 1991.4892; arf, 7 specimens (øsm-10059-25760, 25854 and 25854); mno, 1 specimen. the species is rare. remarks. it differs from the preceding species by having an umbilicus and almost invisible folds on the columella. fig. 164. unitas sp. 2. mguh 31136 (ex gm 1991.4892), height 1.4 mm, width 1.0 mm. genus admetula cossmann 1889 type species. buccinum evulsum solander 1766. admetula rosenkrantzi schnetler & petit 2006 (fig. 165) 2006 admetula rosenkrantzi schnetler & petit, p. 104, figs 6a, b. type material. holotype mguh 27351. additional material. gm 1977.1375, 1 adult specimen. the species is very rare. fig. 165. admetula rosenkrantzi schnetler & petit 2006. mguh 31112 (ex rkz 11), height 7.3 mm, width 5.1 mm. admetula faksensis schnetler & petit 2006 (fig. 166) 2006 admetula faksensis schnetler & petit, p. 104, figs 7a, b. type material. holotype mguh 27356. additional material. arf, 1 juvenile specimen (øsm10059-25855); arf, 1 specimen (øsm-10059-25851); arf, 2 juvenile specimens (øsm-10059-25004 and 25111); isl, 2 fragmented specimens. the species is rare. 102102 fig. 166. admetula faksensis schnetler & petit 2006. mguh 27356 (ex rkz 12), height 10.0 mm, width 6.8 mm. genus semitriton cossmann 1903 type species. plesiotriton dennanti tate 1898. semitriton biplicatus (ravn 1902b) (fig. 167) 1902 tritonium biplicatum (m.u.h.) n. sp. – ravn, p. 228 (24), plate 2, figs 11–13. 1933 admete (?) biplicata (ravn) – ravn, p. 68, plate 1, figs 12a, b. (partim, non plate 1, figs 12a, b = unitas aliceae). 2006 semitriton biplicatus (ravn 1902b) – schnetler & petit, p. 105, figs 5, 8, 17a, b. type material. lectotype mguh 108 (= ravn 1902b, plate 2, fig. 11); mguh 109 (= ravn 1902b, plate 2, fig. 12); mguh 110 (= ravn 1902b, plate 2, fig. 13). additional material. mguh 27353; isl, 10 juvenile specimens; afr, 24 juvenile specimens (øsm-1005925849). the species is rather common. fig. 167. semitriton biplicatus (ravn 1902b) mguh 27352, height 9.8 mm, width 2.8 mm. genus tatara fleming 1950 type species. cymatium pahiense marshall & murdoch 1921. tatara danica schnetler & petit 2006 (fig. 168) 2006 tatara danica schnetler & petit, p. 106, figs 9, 10a, b, 18, 19a, b. type material. holotype mguh 27354, paratypes mguh 27355, mguh 27357, mguh 27358. additional material. gm1977.1373 and 1977.1374, 2 specimens; isl, 17 juvenile specimens; arf, 1 adult and 2 juvenile specimens (øsm-10059-21429 and 25081). the species is rather common. 103 fig. 168. tatara danica schnetler & petit 2006. a, b: mguh 31137 (ex rkz 9), height 5.8 mm, width 3.1 mm. c, d: mguh 31138 (ex rkz 10), height 5.3 mm, width 2.8 mm. clade heterobranchia informal group lower heterobranchia superfamily acteonoidea d’orbigny 1843 family acteonidae d’orbigny 1843 subfamily acteoninae d’orbigny 1843 genus acteon montfort 1810 type species. acteon tornatilis linnaeus 1758. acteon sp. (fig. 169) additional material. mno, 4 specimens. the species is rare. fig. 169. acteon sp. a, b: mguh 31139 (ex rkz 253), eight 1.9 mm, width 1.3 mm. c, d: mguh 31140 (ex rkz 254), height 1.4 mm, width 1.0 mm. genus rictaxis dall 1871 type species. rictaxis punctocaelatus (carpenter 1864). rictaxis? selandica (ravn 1933) (fig. 170) 1933 odostomia? selandica ravn, p. 40, plate 3, figs 13a, b. type material. holotype mguh 3146. 104104 additional material. rkz 255, rkz 256. the species is common. remarks. ravn (1933, p. 41) questioned the assignment to odostomia. kollmann & peel (1983, p. 105, fig. 243; p. 106, fig. 244) illustrated two related species from nuussuaq, sub nomine new genus cf. rictaxis dall, species 1 and 2. the faxe species has a similar protoconch, apex and spiral ornament as cf. rictaxis sp. 2, but the columellar plait is less distinct. on drawing rkz 257, however, the columellar plait could be observed on a defective specimen and it matches the cf. rictaxis sp. 2 very well. for these reasons we tentatively assign the faxe species to rictaxis. fig. 170. rictaxis? selandica (ravn 1933).a, b: mguh 31141 (ex rkz 256b), height 3.4 mm, width 2.2 mm. c: mguh 31142 (ex rkz 257c), height 2.9 mm, width 1.5 mm. acteonoidea gen. et sp. indet. (fig. 171) additional material. gm (ex rkz 153); arf, 21 specimens (øsm-10059-25093); isl, 2 specimens; ebp, 2 specimens. the species is rather common. remarks. the genus name palaeocrenilabium was indicated by rosenkrantz on a drawing, but never published. fig. 171. acteonoidea gen. et sp. indet. a, b: mguh 31143 (ex rkz 231b), height 4.0 mm, width 1.5 mm. c, d: mguh 31144 (ex rkz 231a), diameter of protoconch 0.7 mm. superfamily rissoelloidea gray 1850 family rissoellidae gray 1850 genus rissoella gray 1847a type species. rissoa? glaber alder = rissoella glaber (err. pro glabra) j.e. gray 1847a; = rissoa? diaphana alder 1848; = rissoa albella (alder 1844). rissoella? sp. (fig. 172) material. only the illustrated specimen is known. 105 fig. 172. rissoella? sp. mguh 31145 (ex rkz 146), height 2.2 mm, width 1.4 mm. superfamily architectonicoidea gray 1850 family architectonicidae gray 1850 genus pseudotorinia sacco 1892 type species. solarium obtusum bronn 1831. pseudotorinia faxense (ravn 1933) (fig. 173) 1933 solarium faxense ravn, p. 34, plate 3, figs 1a–c, 2a–c. type material. holotype mguh 3135. material. the species is very common. fig. 173. pseudotorinia faxense (ravn 1933). mguh 3134, height 2.9 mm, width 5.3 mm, diameter of umbilicus 2.3 mm. genus nipteraxis cossmann 1916 type species. solarium plicatum lamarck 1804. nipteraxis poulseni (ravn 1933) (fig. 174) 1933 solarium poulseni ravn, p. 35, plate 3, figs 5a–c. type material. holotype mguh 3138. additional material. arf, 7 specimens. the species is rare. fig. 174. nipteraxis poulseni (ravn 1933). mguh 3138, height 2.6 mm, width 6.0 mm, diameter of umbilicus 3.5 mm. genus pseudomalaxis fischer 1885 type species. bifrontia zanclea philippi 1844. pseudomalaxis sp. (fig. 175) material. the species is rather common, but almost all specimens are very small. 106106 fig. 175. pseudomalaxis sp. øsm-10059-25127 (ex arf), height 1.6 mm, width 4.2 mm. photo: leif rasmussen, faxe, denmark. family orbitestellidae iredale 1917 genus orbitestella iredale 1917 type species. orbitestella bastowi gatliff 1906 orbitestella sp. (fig. 176) material. only the illustrated specimen is known. fig. 176. orbitestella sp. mguh 31147 (ex rkz 101), height 1.8 mm, width 2.1 mm. family amphitomariidae bandel 1996 genus neamphitomaria bandel 1988 type species. pseudomalaxis stantoni sohl 1960. neamphitomaria sp. 1 (fig. 177) additional material. one specimen, rkz 126. remarks. rosenkrantz suggested in his notes that these specimens belonged to the genus omalaxis deshayes 1832. fig. 177. neamphitomaria sp. 1. mguh 31148 (ex rkz 125), height 0.8 mm, width 2.8 mm. neamphitomaria sp. 2 (fig. 178) material. only the illustrated specimen is known. remarks. the species has a keel on the highest part of the whorl and another keel at the transition to the base. the aperture is ovate. fig. 178. neamphitomaria sp. 2. mguh 31149 (ex rkz 113), height 0.6 mm, width 1.7 mm. 107 neamphitomaria sp. 3 (fig. 179) additional material. mno, 1 specimen. the species is very rare. remarks. the species has no keel at the periphery and a higher apex. the aperture is subcircular. b c a fig. 179. neamphitomaria sp. 3. mguh 31150 (ex rkz 124), height 1.2 mm, width 2.6 mm. superfamily mathildoidea dall 1889 family mathildidae dall 1889 genus gegania jeffreys 1884 type species. gegania pinguis jeffreys 1884. gegania rosenkrantzi (ravn 1933) (fig. 180) 1933 mathildia rosenkrantzi ravn, p. 44, plate 4, figs 7a, b. 1933 basilissa? tricincta ravn, p. 32, plate 3, figs 4a, b. type material. holotype mguh 3154. material. the species is common. remarks. ravn (1933) established the species basilissa? tricincta, based on a rather badly preserved specimen. this specimen is here refered to gegania rosenkrantzi, in accordance with the suggestion by rosenkrantz. fig. 180. gegania rosenkrantzi (ravn 1933). a, b: mguh 31151 (ex rkz 85a). a: height 1.6 mm, width 1.3 mm. b: diameter 1.4 mm. c: mguh 31152 (ex rkz 85c), height 5.3 mm, width 3.4 mm. d: mguh 31153 (ex rkz 85b), height 7.1 mm, width 3.9 mm. genus mathilda semper 1865 type species. cerithium fimbriatum michelotti 1847. mathilda unicarinata (ravn 1933) (fig. 181) 1933 mathildia unicarinata ravn, p. 43, plate 4, figs 4a, b. type material. holotype mguh 3151. material. the species is rather common. remarks. the species is characterised by the distinct carina, situated near the adapical suture. there are three weak spirals over the carina and one below. the number of axial ribs is 60–70. 108108 fig. 181. mathilda unicarinata (ravn 1933). a: mguh 31154 (ex rkz 93a), height 1.8, width 1.3 mm. b: mguh 31155 (ex rkz 93b), height 5.7 mm, width 2.6 mm. mathilda sp. 1 (fig. 182) material. only the illustrated specimens are known. remarks. the species has two almost equal spiral ribs and four spiral ribs adapically, resulting in an almost reticulate sculpture. the number of axial ribs is 35–45. a b c d e f fig. 182. mathilda sp. 1. a: mguh 31156 (ex rkz 89a), height 2.2 mm, width 1.1 mm. b: mguh 31157 (ex rkz 90c), height 3.8 mm, width 2.0 mm. c: mguh 31158 (ex rkz 90b), height 2.2 mm, width 1.3 mm. d: mguh 31159 (ex rkz 89b), height 3.5 mm, width 1.4 mm. e: mguh 31160 (ex rkz 91), height 2.6 mm, width 1.1 mm. f: mguh 31161 (ex rkz 90a), height 5.3 mm, width 2.3 mm. mathilda sp. 2 (fig. 183) additional material. isl, 3 specimens. the species is very rare. remarks. the species has four spiral ribs, which on the last whorl are almost equal in strength. they are distinct, also on the spirals, and their number is 40–45. 109 fig. 183. mathilda sp. 2. mguh 31162 (ex rkz 92a), height 9.5 mm, width 3.9 mm, height of protoconch 1.7 mm. genus acrocoelum cossmann 1888 type species. mathilda bouryi cossmann 1888. acrocoelum? sp. 1 (fig. 184) additional material. mno, 2 specimens. the species is very rare. remarks. this species has four spirals and about 30 axial ribs, which cause a reticulate pattern. fig. 184. acrocoelum? sp. 1. mguh 31163 (ex rkz 86), height 1.9 mm, width 0.8 mm. acrocoelum sp. 2 (fig. 185) additional material. isl, 1 specimen; mno, 2 specimens. the species is very rare. remarks. this species has relatively high whorls with eight spiral bands and very weak axial ribs. mathilda lemchei ravn 1939 (p. 68, plate 2, figs 16a, b) has only six spiral bands, but resembles the species in general outline. fig. 185. acrocoelum sp. 2. øsm-10059-25098 (ex arf), height 6.8 mm, width 2.0 mm. genus clathrobaculus cossmann 1912 type species. cerithium ziczac eudes-deslongchamps 1842. clathrobaculus? sp. 1 (fig. 186) additional material. isl, 1 specimen. the species is very rare. remarks. this species has six spiral bands, of which the abapical two are the strongest. 110110 fig. 186. clathrobaculus? sp. 1. mguh 31165 (ex rkz 95), height 4.7 mm, width 1.8 mm. clathrobaculus? sp. 2 (fig. 187) additional material. isl, 10 specimens. the species is rare. remarks. this species has four spiral bands of almost equal strength, with secondary weaker spirals inserted on the younger whorls. the protoconch is unknown. fig. 187. clathrobaculus? sp. 2. mguh 31166 (ex rkz 185), height 5.8 mm, width 1.4 mm. superfamily pyramidelloidea gray 1840 family pyramidellidae gray 1840 genus odostomia fleming 1813 type species. odostomia plicata (montagu 1803). odostomia sp. 1 (fig. 188) material. only the illustrated specimen is known. remarks. the species has a rather high apex and very convex whorls. the columellar tooth is situated higher than the middle of the columella. there is no umbilicus. fig. 188. odostomia sp. 1. mguh 31167 (ex rkz 134), height 2.8 mm, width 1.2 mm. odostomia sp. 2 (fig. 189) additional material. arf, 58 specimens (øsm-1005925045, 25071 and 25821). the species is common. remarks. the species is smaller than the preceding species and has less convex whorls, separated by a less deep suture. furthermore, it has a distinct umbilicus. 111 fig. 189. odostomia sp. 2. mguh 31168 (ex rkz 152), height 1.5 mm, width 0.7 mm, height of protoconch 0.2 mm. family amathinidae ponder 1987 genus leucotina a. adams 1860 type species. leucotina dianae a. adams in h. adams & a. adams 1854. leucotina sp. (fig. 190) additional material. isl, 4 specimens; mno, 2 specimens. the species is rare. fig. 190. leucotina sp. mguh 31169 (ex rkz 147), height 3.5 mm, width 2.2 mm, height of protoconch 0.7 mm. subfamily syrnolinae saurin 1958 genus puposyrnola cossmann 1921 type species. auricula acicula lamarck 1804. puposyrnola sp. (fig. 191) material. the illustrated specimen is the only known example. fig. 191. puposyrnola sp. mguh 31170 (ex rkz 142), height 3.1 mm, width 0.8 mm. 112112 clade opisthobranchia informal group opisthobranchia clade cephalaspidea superfamily philinoidea gray 1850 (1815) family cylichnidae h. & a. adams 1854 genus acteocina gray 1847a type species. bulla voluta quoy & gaimard 1833. acteocina sp. (fig. 192) additional material. rkz 259; arf, 7 specimens; isl, 1 specimen. the species is rare. fig. 192. acteocina sp. mguh 31171 (ex rkz 259), height 1.9 mm, width 1.1 mm. incertae sedis incertae sedis sp. 1 (fig. 193) additional material. gm 1991.4246, 2 specimens. the species is very rare. description: the species is slender with almost flat whorls, separated by a rather distinct suture. there is a rounded carina a little above the middle of the whorl. the shell is smooth, without spirals and axial sculpture, and the aperture is small and subcircular with a concave columella and a very short canal. fig. 193. incertae sedis sp. 1. mguh 31172 (ex gm 1991.4246), height 2.8 mm, width 0.9 mm. incertae sedis sp. 2 (fig. 194) additional material. gm 1991.4247a, 1 specimen. the species is very rare. description: the species has convex whorls, separated by a deep suture, with no visible spiral ornament. on the first teleoconch whorls there are weak opisthocline axial ribs but they fade out on the younger whorls. the aperture is subelliptical. fig. 194. incertae sedis sp. 2. mguh 31173 (ex gm 1991.4247a), height 2.3 mm, width 0.9 mm. 113 references adams, a. 1860: on some new genera and species of mollusca from japan. annals and magazine of natural history series 3, 5, 299– 303, 405–413. adams a. 1861: on some new species of eulima, leiostraca and cerithiopsis from japan. annals and magazine of natural history series 3, 7, 125–131. adams, h. 1860: description of a new genus of shells from the collection of hugh cuming, esq. proceedings of the zoological society of london 28, 272–273. adams, h. & adams, a. 1853: the genera of recent mollusca; arranged according to their organization 1, 1–256. london: johan van voorst. adams, h. & adams, a. 1854: the genera of recent mollusca; arranged according to their organization 1, 257–484. london: johan van voorst. andreae, a. 1887: die glossophoren des terrain à chailles der pfirt. abhandlungen der geologischen specialkarte von elsaß-lothringen 4, 1–45. baluk, w. 1975. lower tortonian gastropods from korytnica, poland, part i. palaeontologia polonica 32, 186 pp. bandel, k. 1988: repräsentieren die euomphaloidea eine natürliche einheit der gastropoden? mitteilungen aus dem geologischpaläontologischen institut der universität hamburg 67, 1–33. bandel. k. 1998: scissurellidae als modell für die variationsbreite einer natürlichen einheit der schlitzbandschnecken (mollusca, archaeogastropoda). mitteilungen aus dem geologisch-paläontologischen institut der universität hamburg 81, 1–120. bandel, k. 2006. families of the cerithioidea and related superfamilies (palaeo-caeogastropoda; mollusca) from the triassic to the recent characterized by protoconch morphology including the description of new taxa.  freiberger forschungshefte, geowissenschaften c511, 59–138. bandel, k. 2010: relationships of the triassic eucycloidea koken, 1897 (mollusca, gastropoda) to modern genera such as pagodatrochus, calliotropis and euchelus, based on morphology of the early shell. bulletin of geosciences 85, 435–486. bandel, k. & kowalke, t. 1997: systematic value of the larval shell of fossil and modern vanikoridae, pickworthiidae and the genus fossarus (caenogastropoda, mollusca). berliner geowissenschaftliche abhandlungen e25, 3–29. berlin. bartsch, p. 1911: the recent and fossil mollusks of the genus cerithiopsis from the west coast of america. proceedings of the united states national museum 40, 327–367. bayan, j.f. 1873: sur le travail de récensement des espèces publiées et sur quelques synonymies. bulletin de la société géologique de france 1, 235 pp. bellardi, l. 1873 : i molluschi dei terreni terziarii del piemonte e della liguria1: cephalopoda, pteropoda, heteropoda, gasteropoda (muricidae e tritonidae). memorie della reale accademia delle scienze di torino, series 2, 27, 33–294 (reprint 264 pp.). bellardi, a. 1886: i molluschi dei terreni terziarii del piemonte e della liguria. parte iv fasciolariidae e turbinellidae. memorie della reale accademia delle scienze di torino, series 2, 37, 1–62. bernecker, m. & weidlich, o. 1990: the danian (paleocene) coral limestone of fakse, denmark: a model for ancient aphotic, azooxanthellate coral mounds. facies 22, 103–138. bernecker, m. & weidlich, o. 2005: azooxanthellate corals in the late maastrichtian – early paleocene of the danish basin: bryozoan and coral mounds in a boreal shelf setting. in: freiwald, a. & roberts, j.m. 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& cairns, s.d. 1992: revision of the northeast atlantic and mediterranean stylasteridae (cnidaria: hydrozoa). mémoires du muséum national d'histoire naturelle. série a, zoologie 153, 1–136. http://nmita.iowa.uiowa.edu/database/mollusc/mollusclifestyles.htm http://nmita.iowa.uiowa.edu/database/mollusc/mollusclifestyles.htm 118118 de nationale geologiske undersøgelser for danmark og grønland (geus) geological survey of denmark and greenland øster voldgade 10, dk-1350 copenhagen k denmark the series geological survey of denmark and greenland bulletin started in 2003 and replaced the two former bulletin series of the survey, viz. geology of greenland survey bulletin and geology of denmark survey bulletin. the twenty-one volumes published since 1997 in those two series are listed on the following pages. the present series, together with geological survey of denmark and greenland map series, now form the peer-reviewed scientific series of the survey. geological survey of denmark and greenland bulletin 1 the jurassic of denmark and greenland, 948 pp. (28 articles), 2003. edited by j.r. ineson & f. surlyk. 500.00 2 fish otoliths from the paleocene of denmark, 94 pp., 2003. by w. schwarzhans. 100.00 3 late quaternary environmental changes recorded in the danish marine molluscan faunas, 268 pp., 2004. by k.s. petersen. 200.00 4 review of survey activities 2003, 100 pp. (24 articles), 2004. edited by m. sønderholm & a.k. higgins. 180.00 5 the jurassic of north-east greenland, 112 pp. (7 articles), 2004. edited by l. stemmerik & s. stouge. 160.00 6 east greenland caledonides: stratigraphy, structure and geochronology, 93 pp. (6 articles), 2004. edited by a.k. higgins & f. kalsbeek. 160.00 7 review of survey activities 2004, 80 pp. (19 articles), 2005. edited by m. sønderholm & a.k. higgins. 180.00 8 structural analysis of the rubjerg knude glaciotectonic complex, vendsyssel, northern denmark, 192 pp., 2005. by s.a.s. pedersen. 300.00 9 scientific results from the deepened lopra-1 borehole, faroe islands, 156 pp. (11 articles), 2006. edited by j.a. chalmers & r. waagstein. 240.00 10 review of survey activities 2005, 68 pp. (15 articles), 2006. edited by m. sønderholm & a.k. higgins. 180.00 11 precambrian crustal evolution and cretaceous–palaeogene faulting in west greenland, 204 pp. (12 articles), 2006. edited by a.a. garde & f. kalsbeek. 240.00 12 lithostratigraphy of the palaeogene – lower neogene succession of the danish north sea, 77 pp., 2007. by p. schiøler, j. andsbjerg, o.r. clausen, g. dam, k. dybkjær, l. hamberg, c. heilmann-clausen, e.p. johannessen, l.e. kristensen, i. prince & j.a. rasmussen. 240.00 13 review of survey activities 2006, 76 pp. (17 articles), 2007. edited by m. sønderholm & a.k. higgins. 180.00 14 quaternary glaciation history and glaciology of jakobshavn isbræ and the disko bugt region, west greenland: a review, 78 pp., 2007. by a. weidick & o. bennike. 200.00 15 review of survey activities 2007, 96 pp. (22 articles), 2008. edited by o. bennike & a.k. higgins. 200.00 16 evaluation of the quality, thermal maturity and distribution of potential source rocks in the danish part of the norwegian–danish basin, 66 pp., 2008. by h.i. petersen, l.h. nielsen, j.a. bojesen-koefoed, a. mathiesen, l. kristensen & f. dalhoff. 200.00 17 review of survey activities 2008, 84 pp. (19 articles), 2009. edited by o. bennike, a.a. garde & w.s. watt. 200.00 18 greenland from archaean to quaternary. descriptive text to the 1995 geological map of greenland, 1:2 500 000. 2nd edition, 126 pp., 2009. by n. henriksen, a.k. higgins, f. kalsbeek & t.c.r. pulvertaft. 280.00 19 lithostratigraphy of the cretaceous–paleocene nuussuaq group, nuussuaq basin, west greenland, 171 pp., 2009. by g. dam, g.k. pedersen, m. sønderholm, h.h. midtgaard, l.m. larsen, h. nøhr-hansen & a.k. pedersen. 300.00 20 review of survey activities 2009, 106 pp. (23 articles), 2010. edited by o. bennike, a.a. garde & w.s. watt. 220.00 21 exploration history and place names of northern east greenland, 368 pp., 2010. by a.k. higgins. 200.00 22 lithostratigraphy of the upper oligocene – miocene succession of denmark, 92 pp., 2010. by e.s. rasmussen, k. dybkjær & s. piasecki. 240.00 23 review of survey activities 2010, 84 pp. (19 articles), 2011. edited by o. bennike, a.a. garde & w.s. watt. 200.00 24 the east greenland rifted volcanic margin, 96 pp., 2011. by c.k. brooks. 200.00 25 upper cretaceous chalk facies and depositional history recorded in the mona-1 core, mona ridge, 119 danish north sea. 2011. by k. anderskouv & f. surlyk. 200.00 26 review of survey activities 2011, 88 pp. (21 articles), 2012. edited by o. bennike, a.a. garde & w.s. watt. 200.00 27 neoglacial and historical glacier changes around kangersuneq fjord in southern west greenland, 68 pp., 2012. by a. weidick, o. bennike, m. citterio & n. nørgaard-pedersen. 200.00 28 review of survey activities 2012, 76 pp. (17 articles), 2013. edited by o. bennike, a.a. garde & w.s. watt. 200.00 29 tectono-magmatic evolution of the younger gardar southern rift, south greenland, 124 pp., 2013.  by b.g.j. upton. 240.00 30 stratigraphic landscape analysis, thermochronology and the episodic development of elevated, passive continental margins, 150 pp., 2014. by paul f. green, karna lidmar-bergström, peter japsen, johan m. bonow and james a. chalmers. 250.00 31 review of survey activities 2013, 98 pp., 2014. edited by o. bennike, a.a. garde & w.s. watt 200.00 32 a catalogue of danian gastropods from the baunekule facies, faxe formation, denmark, 117 pp., 2014 by b.w. lauridsen & k.i. schnetler. 240.00 geological survey of denmark and greenland map series 1 explanatory notes to the geological map of greenland, 1:500 000, humboldt gletscher, sheet 6, 48 pp. + map, 2004. by p.r. dawes. 280.00 2 explanatory notes to the geological map of greenland, 1:500 000, thule, sheet 5 (1991), 97 pp. + map, 2006. by p.r. dawes. 300.00 3 explanatory notes to the geological map of greenland, 1:100 000, ussuit 67 v.2 nord, 40 pp. + map, 2007. by j.a.m. van gool & m. marker. 280.00 4 descriptive text to the geological map of greenland, 1:500 000, dove bugt, sheet 10, 32 pp. + map, 2009. by n. henriksen & a.k. higgins. 240.00 5 descriptive text to the geological map of greenland, 1:100 000, kangaatsiaq 68 v.1 syd and ikamiut 68 v.1 nord, 41 pp. + 2 maps, 2010. by a.a. garde & j.a. hollis. 280.00 geology of greenland survey bulletin (173–191; discontinued) 173 cambrian shelf stratigraphy of north greenland, 120 pp., 1997. by j.r. ineson & j.s. peel 250.00 174 the proterozoic thule supergorup, greenland and canada: history, lithostratigraphy and development, 150 pp., 1997. by p.r. dawes 300.00 175 stratigraphy of the neill klinter group; a lower – lower middle jurassic tidal embayment succession, jameson land, east greenland, 80 pp., 1998. by g. dam & f. surlyk. 250.00 176 review of greenland activities 1996, 112 pp. (18 articles), 1997. edited by a.k. higgins & j.r. ineson. 200.00 177 accretion and evolution of an archaean high-grade grey gneiss – amphibolite complex: the fiskefjord area, southern west greenland, 115 pp., 1997. by a.a. garde. 200.00 178 lithostratigraphy, sedimentary evolution and sequence stratigraphy of the upper proterozoic lyell land group (eleonore bay supergroup) of east and north-east greenland, 60 pp., 1997. by h. tirsgaard & m. sønderholm. 200.00 179 the citronen fjord massive sulphide deposit, peary land, north greenland: discovery, stratigraphy, mineralization and structural setting, 40 pp., 1998. by f.w. van der stijl & g.z. mosher. 200.00 180 review of greenland activities 1997, 176 pp. (26 articles), 1998. edited by a.k. higgins & w.s. watt. 200.00 181 precambrian geology of the disko bugt region, west greenland, 179 pp. (15 articles), 1999. edited by f. kalsbeek. 240.00 182 vertebrate remains from upper silurian – lower devonian beds of hall land, north greenland, 80 pp., 1999. by h. blom. 120.00 183 review of greenland activities 1998, 81 pp. (10 articles), 1999. edited by a.k. higgins & w.s. watt. 200.00 184 collected research papers: palaeontology, geochronology, geochemistry, 62 pp. (6 articles), 1999. 150.00 185 greenland from archaean to quaternary. descriptive text to the geological map of greenland, 1:2 500 000, 93 pp., 2000. by n. henriksen, a.k. higgins, f. kalsbeek & t.c.r. pulvertaft. 225.00 120120 186 review of greenland activities 1999, 105 pp. (13 articles), 2000. edited by p.r. dawes & a.k. higgins. 225.00 187 palynology and deposition in the wandel sea basin, eastern north greenland, 101 pp. (6 articles), 2000. edited by l. stemmerik. 160.00 188 the structure of the cretaceous–palaeogene sedimentary-volcanic area of svartenhuk halvø, central west greenland, 40 pp., 2000. by j. gutzon larsen & t.c.r. pulvertaft. 130.00 189 review of greenland activities 2000, 131 pp. (17 articles), 2001. edited by a.k. higgins & k. secher. 160.00 190 the ilímaussaq alkaline complex, south greenland: status of mineralogical research with new results, 167 pp. (19 articles), 2001. edited by h. sørensen. 160.00 191 review of greenland activities 2001, 161 pp. (20 articles), 2002. edited by a.k. higgins, k. secher & m. sønderholm. 200.00 geology of denmark survey bulletin (36–37; discontinued) 36 petroleum potential and depositional environments of middle jurassic coals and non-marine deposits, danish central graben, with special reference to the søgne basin, 78 pp., 1998. by h.i. petersen, j. andsbjerg, j.a. bojesen-koefoed, h.p. nytoft & p. rosenberg. 250.00 37 the selandian (paleocene) mollusc fauna from copenhagen, denmark: the poul harder 1920 collection, 85 pp., 2001. by k.i. schnetler. 150.00 prices are in danish kroner exclusive of local taxes, postage and handling note that information on the publications of the former geological survey of denmark and the former geological survey of greenland (amalgamated in 1995 to form the present geological survey of denmark and greenland) can be found on www.geus.dk http://www.geus.dk a catalogue of danian gastropods from the baunekule facies, faxe formation, denmark frontispiece contents abstract introduction material previous work on gastropods from the baunekule facies geological setting notes on the locality and the baunekule facies remarks on the fauna preservation number of taxa stratigraphic ranges of the genera comparison with other paleocene gastropod faunas palaeoecology gastropods on modern cold-water coral mounds concluding remarks and future studies acknowledgements systematic palaeontology abbreviations and classification references geological survey of denmark and greenland bulletin 36 geological survey of denmark and greenland bulletin 36· 2016 cretaceous and cenozoic dinoflagellate cysts and other palynomorphs from the western and eastern margins of the labrador–baffin seaway robert a. fensome, henrik nøhr-han sen & graham l. williams geological survey of denmark and greenland ministry of energy, utilities and climate bulletin36.qxp_bulletin 36 19/12/16 13.39 side 1 geological survey of denmark and greenland bulletin 36 keywords biostratigraphy, cretaceous, cenozoic, dinocysts, systematics, baffin margin, labrador margin, mesozoic, offshore west greenland cover illustration selected palynomorphs from the subsurface of the labrador–baffin seaway; for details, see plates 1, 3, 4, 6, 7, 12, 15, 17, 18. top left to bottom right: dinoflagellates – chiropteridium gilbertii sp. nov., chatangiella tripartite, cleistosphaeridium palmatum sp. nov., diphyes brevispinum, ginginodinium? flexidentatum sp. nov., adnatosphaeridium vittatum, piladinium columna, thalassiphora pelagica, eocladopyxis peniculata; miospores – aquilapollenites quadrilobus, cicatricososporites eocenicus, baculatisporites crenulatus sp. nov. frontispiece: facing page the canadian icebreaker amundsen off beechey island in the canadian arctic in the autumn of 2013. geological research in the arctic is entirely dependent on such professional logistic support, whether by sea or air. photo: kate jarrett (gsc). chief editor of this series: adam a. garde scientific editor of this volume: jon r. ineson editorial secretary: jane holst referees: martin pearce (uk) and james b. riding (uk) illustrators: jette halskov (geus) and bill macmillan (gsca) digital photographic work: benny m. schark graphic production: annabeth andersen printers: rosendahls · schultz grafisk a/s, albertslund, denmark manuscript received: 21 january 2015 final version approved: 17 march 2016 printed: 19 december 2016 issn (print) 1604-8156 issn (online) 1904-4666 isbn (print) 978-87-7871-443-5 isbn (online) 978-87-7871-444-2 citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bulletin 36, 143 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk or at www.geus.dk/publications/bull © de nationale geologiske undersøgelser for danmark og grønland (geus), 2016 for the full text of the geus copyright clause, please refer to www.geus.dk/publications/bull bulletin36.qxp_bulletin 36 19/12/16 13.39 side 2 bulletin36.qxp_bulletin 36 19/12/16 13.39 side 3 4 contents abstract . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . methodology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . systematics – general . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . systematics – dinoflagellates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . division dinoflagellata class dinophyceae genus achilleodinium eaton 1976 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . achilleodinium biformoides (eisenack 1954) eaton 1976 genus adnatosphaeridium williams & downie 1966a . . . . . . . . . . . . . . . . . . . . . . . . adnatosphaeridium vittatum williams & downie 1966a genus alisocysta stover & evitt 1978 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . alisocysta circumtabulata (drugg 1967) stover & evitt 1978 alisocysta margarita (harland 1979a) harland 1979a genus alterbidinium lentin & williams 1985 emend. nov. . . . . . . . . . . . . . . . . . . . . alterbidinium acutulum (wilson 1967a) lentin & williams 1985 alterbidinium biaperturum (mcintyre 1975) comb. nov. alterbidinium? bicellulum (islam 1983a) lentin & williams 1985 alterbidinium ioannidesii pearce 2010 alterbidinium varium kirsch 1991 genus apectodinium (costa & downie 1976) lentin & williams 1977a emend. williams, damassa, fensome & guerstein in fensome et al. 2009 . . . . . . . . . . . . . . . . . . . apectodinium homomorphum (deflandre & cookson 1955) lentin & williams 1977a apectodinium parvum (alberti 1961) lentin & williams 1977a apectodinium quinquelatum (williams & downie 1966b) costa & downie 1979 genus aptea eisenack 1958 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . aptea polymorpha eisenack 1958 genus apteodinium eisenack 1958 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . apteodinium australiense (deflandre & cookson 1955) williams 1978 apteodinium spiridoides benedek 1972 genus areoligera lejeune-carpentier 1938 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . areoligera circumsenonensis fensome et al. 2009 areoligera gippingensis jolley 1992 genus areosphaeridium eaton 1971 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . areosphaeridium diktyoplokum (klumpp 1953) eaton 1971 genus atopodinium drugg 1978 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . atopodinium cf. haromense thomas & cox 1988 genus axiodinium williams et al. in fensome et al. 2009 . . . . . . . . . . . . . . . . . . . . . . axiodinium augustum (harland 1979b) williams et al. 2015 genus batiacasphaera drugg 1970 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . batiacasphaera micropapillata stover 1977 genus batioladinium brideaux 1975 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . batioladinium jaegeri (alberti 1961) brideaux 1975 genus callaiosphaeridium davey & williams 1966a . . . . . . . . . . . . . . . . . . . . . . . . . . callaiosphaeridium asymmetricum (deflandre & courteville 1939) davey & williams 1966a 13 15 17 23 23 23 23 24 24 26 27 27 28 29 29 29 29 30 30 bulletin36.qxp_bulletin 36 19/12/16 13.39 side 4 genus cannosphaeropsis wetzel 1933a . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . cannosphaeropsis passio de verteuil & norris 1996 genus cerebrocysta bujak in bujak et al. 1980 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . cerebrocysta bartonensis bujak in bujak et al. 1980 cerebrocysta magna bujak 1994 . genus cerodinium vozzhennikova 1963 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . cerodinium diebelii (alberti 1959) lentin & williams 1987 cerodinium glabrum (gocht 1969) fensome et al. 2009 cerodinium kangiliense nøhr-hansen & heilmann-clausen 2001 cerodinium speciosum (alberti 1959) lentin & williams 1987 cerodinium striatum (drugg 1967) lentin & williams 1987 genus charlesdowniea lentin & vozzhennikova 1989 emend. williams et al. 2015 . . . charlesdowniea coleothrypta (williams & downie 1966b) lentin & vozzhennikova 1989 genus chatangiella vozzhennikova 1967 emend. nov. . . . . . . . . . . . . . . . . . . . . . . . . chatangiella decorosa (mcintyre 1975) lentin & williams 1976 chatangiella madura lentin & williams 1976 chatangiella tripartita (cookson & eisenack 1960a) lentin & williams 1976 genus chiropteridium gocht 1960 emend. nov. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . chiropteridium galea (maier 1959) sarjeant 1983 chiropteridium gilbertii sp. nov. genus chlamydophorella cookson & eisenack 1958 . . . . . . . . . . . . . . . . . . . . . . . . . . chlamydophorella nyei cookson & eisenack 1958 chlamydophorella cf. nyei cookson & eisenack 1958 genus chytroeisphaeridia (sarjeant 1962) downie & sarjeant 1965 . . . . . . . . . . . . . . chytroeisphaeridia hadra sp. nov. genus cleistosphaeridium davey et al. 1966 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . cleistosphaeridium diversispinosum davey et al. 1966 cleistosphaeridium elegantulum sp. nov. cleistosphaeridium palmatum sp. nov. cleistosphaeridium polypetellum (islam 1983b) stover & williams 1995 genus cordosphaeridium eisenack 1963a . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . cordosphaeridium cantharellus (brosius 1963) gocht 1969 cordosphaeridium delimurum fensome et al. 2009 cordosphaeridium fibrospinosum davey & williams 1966a cordosphaeridium funiculatum morgenroth 1966a cordosphaeridium gracile (eisenack 1954) davey & williams 1966a cordosphaeridium inodes (klumpp 1953) eisenack 1963a genus cribroperidinium neale & sarjeant 1962 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . cribroperidinium giuseppei (morgenroth 1966a) helenes 1984 genus cyclonephelium deflandre & cookson 1955 . . . . . . . . . . . . . . . . . . . . . . . . . . . cyclonephelium distinctum (deflandre & cookson 1955) jansonius 1986 genus dapsilidinium bujak et al. 1980 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . dapsilidinium pseudocolligerum (stover 1977) bujak et al. 1980 dapsilidinium pseudoinsertum sp. nov. dapsilidinium simplex (white 1842) bujak et al. 1980 5 30 31 31 32 32 34 35 36 36 38 39 39 40 bulletin36.qxp_bulletin 36 19/12/16 13.39 side 5 genus deflandrea eisenack 1938 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . deflandrea borealis sp. nov. deflandrea denticulata alberti 1959 deflandrea galeata (lejeune-carpentier 1942) lentin & williams 1973 deflandrea majae (schiøler 1993) comb. nov. deflandrea oebisfeldensis alberti 1959 deflandrea phosphoritica eisenack 1938 genus dinogymnium evitt et al. 1967 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . dinogymnium longicorne (vozzhennikova 1967) harland 1973 genus diphyes cookson 1965 nom. cons. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . diphyes brevispinum bujak 1994 diphyes colligerum (deflandre & cookson 1955) cookson 1965 diphyes ficusoides islam 1983a genus disphaerogena wetzel 1933a . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . disphaerogena carposphaeropsis wetzel 1933a genus eatonicysta stover & evitt 1978 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . eatonicysta furensis (heilmann-clausen in heilmann-clausen & costa 1989) stover & williams 1995 eatonicysta ursulae (morgenroth 1966a) stover & evitt 1978 genus enneadocysta stover & williams 1995 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . enneadocysta magna fensome et al. 2007 genus eocladopyxis morgenroth 1966a . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . eocladopyxis peniculata morgenroth 1966a genus evittosphaerula manum 1979 emend. damassa 1997 . . . . . . . . . . . . . . . . . . . . evittosphaerula? foraminosa sp. nov. genus fibrocysta stover & evitt 1978 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . fibrocysta bipolaris (cookson & eisenack 1965a) stover & evitt 1978 genus gillinia cookson & eisenack 1960a . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . gillinia hymenophora cookson & eisenack 1960a genus ginginodinium cookson & eisenack 1960a . . . . . . . . . . . . . . . . . . . . . . . . . . . ginginodinium? flexidentatum sp. nov. genus glaphyrocysta stover & evitt 1978 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . glaphyrocysta divaricata (williams & downie 1966a) stover & evitt 1978 glaphyrocysta exuberans (deflandre & cookson 1955 ex eaton 1976) stover & evitt 1978 glaphyrocysta retiintexta cookson 1965 glaphyrocysta texta (bujak 1976) stover & evitt 1978 glaphyrocysta vicina (eaton 1976) stover & evitt 1978 genus habibacysta head et al. 1989 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . habibacysta tectata head et al. 1989 genus hapsocysta davey 1979 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . hapsocysta? benteae nøhr-hansen 1993 genus heteraulacacysta drugg & loeblich jr. 1967 . . . . . . . . . . . . . . . . . . . . . . . . . . . heteraulacacysta porosa bujak in bujak et al. 1980 genus heterosphaeridium cookson & eisenack 1968 . . . . . . . . . . . . . . . . . . . . . . . . . . heterosphaeridium bellii radmacher et al. 2014 heterosphaeridium difficile (manum & cookson 1964) ioannides 1986 genus histiocysta davey 1969a . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . histiocysta palla davey 1969a genus homotryblium davey & williams 1966a . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6 40 42 42 43 43 43 44 44 45 45 45 46 47 47 48 48 49 49 bulletin36.qxp_bulletin 36 19/12/16 13.39 side 6 homotryblium abbreviatum eaton 1976 homotryblium tenuispinosum davey & williams 1966a genus hystrichokolpoma klumpp 1953 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . hystrichokolpoma cinctum klumpp 1953 hystrichokolpoma globulus michoux 1985 genus hystrichosphaeridium deflandre 1937 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . hystrichosphaeridium quadratum sp. nov. hystrichosphaeridium tubiferum (ehrenberg 1838) deflandre 1937 genus hystrichosphaeropsis deflandre 1935 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . hystrichosphaeropsis perforata schiøler 1993 hystrichosphaeropsis quasicribrata (wetzel 1961) gocht 1976 genus hystrichostrogylon agelopoulos 1964 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . hystrichostrogylon digitus sp. nov. genus impagidinium stover & evitt 1978 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . impagidinium victorianum (cookson & eisenack 1965b) stover & evitt 1978 genus impletosphaeridium morgenroth 1966a . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . impletosphaeridium apodastum sp. nov. genus isabelidinium lentin & williams 1977b . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . isabelidinium cooksoniae (alberti 1959) lentin & williams 1977b isabelidinium cretaceum (cookson 1956) lentin & williams 1977b isabelidinium microarmum (mcintyre 1975) lentin & williams 1977b genus kiokansium stover & evitt 1978 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . kiokansium williamsii singh 1983 genus kleithriasphaeridium davey 1974 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . kleithriasphaeridium mantellii (davey & williams 1966a) comb. nov. genus laciniadinium mcintyre 1975 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . laciniadinium arcticum (manum & cookson 1964) lentin & williams 1980 genus lentinia bujak in bujak et al. 1980 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . lentinia serrata bujak in bujak et al. 1980 genus licracysta fensome et al. 2007 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . licracysta corymbus fensome et al. 2007 licracysta? semicirculata (morgenroth 1966b) fensome et al. 2007 genus lingulodinium wall 1967 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . lingulodinium funginum (morgenroth 1966a) islam 1983a lingulodinium machaerophorum (deflandre & cookson 1955) wall 1967 genus nyktericysta bint 1986 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . nyktericysta davisii bint 1986 nyktericysta dictyophora he chengquan et al. 1992 nyktericysta tripenta (bint 1986) fensome et al. 2009 genus odontochitina deflandre 1937 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . odontochitina ancala bint 1986 odontochitina costata alberti 1961 odontochitina porifera cookson 1956 genus oligosphaeridium davey & williams 1966a . . . . . . . . . . . . . . . . . . . . . . . . . . . . oligosphaeridium albertense (pocock 1962) davey & williams 1969 oligosphaeridium pulcherrimum (deflandre & cookson 1955) davey & williams 1966a oligosphaeridium totum brideaux 1971 genus operculodinium wall 1967 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . operculodinium centrocarpum (deflandre & cookson 1955) wall 1967 7 49 50 50 51 51 52 52 53 54 54 55 55 55 56 57 57 58 bulletin36.qxp_bulletin 36 19/12/16 13.39 side 7 genus palaeocystodinium alberti 1961 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . palaeocystodinium bulliforme ioannides 1986 palaeocystodinium golzowense alberti 1961 palaeocystodinium teespinosum fensome et al. 2009 genus palaeohystrichophora deflandre 1935 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . palaeohystrichophora infusorioides deflandre 1935 genus palaeoperidinium deflandre 1934 ex sarjeant 1967 . . . . . . . . . . . . . . . . . . . . . . palaeoperidinium pyrophorum (ehrenberg 1838 ex wetzel 1933b) sarjeant 1967 genus palynodinium gocht 1970 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . palynodinium grallator gocht 1970 genus petalodinium williams et al. 2015 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . petalodinium condylos (williams & downie 1966b) williams et al. 2015 genus phelodinium stover & evitt 1978 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . phelodinium kozlowskii (górka 1963) lindgren 1984 genus phthanoperidinium drugg & loeblich jr. 1967 . . . . . . . . . . . . . . . . . . . . . . . . phthanoperidinium coreoides (benedek 1972) lentin & williams 1976 phthanoperidinium levimurum bujak in bujak et al. 1980 phthanoperidinium multispinum bujak in bujak et al. 1980 phthanoperidinium regale bujak 1994 phthanoperidinium stockmansii (de coninck 1975) lentin & williams 1977a genus piladinium williams et al. 2015 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . piladinium columna (michoux 1988) williams et al. 2015 piladinium edwardsii (wilson 1967b) williams et al. 2015 genus pseudoceratium gocht 1957 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . pseudoceratium sp. genus raphidodinium deflandre 1936 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . raphidodinium fucatum deflandre 1936 genus reticulatosphaera matsuoka 1983 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . reticulatosphaera actinocoronata (benedek 1972) bujak & matsuoka 1986 genus rhombodinium gocht 1955 emend. williams et al. in fensome et al. 2009 . . . . rhombodinium draco gocht 1955 rhombodinium porosum bujak 1979 genus rottnestia cookson & eisenack 1961a . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . rottnestia borussica (eisenack 1954) cookson & eisenack 1961a genus scalenodinium gen. nov. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . scalenodinium scalenum sp. nov. genus schematophora deflandre & cookson 1955 . . . . . . . . . . . . . . . . . . . . . . . . . . . . schematophora speciosa deflandre & cookson 1955 genus senegalinium jain & millepied 1973 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . senegalinium iterlaaense nøhr-hansen & heilmann-clausen 2001 genus senoniasphaera clarke & verdier 1967 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . senoniasphaera inornata (drugg 1970) stover & evitt 1978 senoniasphaera microreticulata brideaux & mcintyre 1975 senoniasphaera rotundata clarke & verdier 1967 genus simplicidinium gen. nov. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . simplicidinium insolitum (eaton 1976) comb. nov. genus sophismatia williams et al. 2015 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . sophismatia tenuivirgula (williams & downie 1966b) williams et al. 2015 genus spinidinium cookson & eisenack 1962 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . spinidinium echinoideum (cookson & eisenack 1960a) lentin & williams 1976 8 59 59 59 60 60 61 61 62 63 63 63 64 64 64 65 65 65 66 67 67 bulletin36.qxp_bulletin 36 19/12/16 13.39 side 8 genus spiniferites mantell 1850 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . spiniferites ovatus matsuoka 1983 spiniferites pseudofurcatus (klumpp 1953) sarjeant 1970 spiniferites scabrosus (clarke & verdier 1967) lentin & williams 1975 genus spongodinium deflandre 1936 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . spongodinium delitiense (ehrenberg 1838) deflandre 1936 spongodinium grossum (manum & cookson 1964) comb. nov. spongodinium obscurum (manum & cookson 1964) comb. nov. genus stichodinium williams et al. 2015 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . stichodinium lineidentatum (deflandre & cookson 1955) williams et al. 2015 genus subtilisphaera jain & millepied 1973 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . subtilisphaera perlucida (alberti 1959) jain & millepied 1973 genus surculosphaeridium davey et al. 1966 emend. nov. . . . . . . . . . . . . . . . . . . . . . . surculosphaeridium convocatum sp. nov. genus talladinium williams, damassa, fensome & guerstein in fensome et al. 2009 . . talladinium? clathratum (eisenack 1938) williams, damassa, fensome & guerstein in fensome et al. 2009 talladinium pellis sp. nov. genus tanyosphaeridium davey & williams 1966a . . . . . . . . . . . . . . . . . . . . . . . . . . . tanyosphaeridium xanthiopyxides (wetzel 1933a ex deflandre 1937) stover & evitt 1978 genus taurodinium gen. nov. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . taurodinium granulatum sp. nov. genus tenua eisenack 1958 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . tenua hystrix eisenack 1958 genus thalassiphora eisenack & gocht 1960 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . thalassiphora delicata williams & downie 1966a thalassiphora fenestrata liengjarern et al. 1980 thalassiphora pelagica (eisenack 1954) eisenack & gocht 1960 genus trichodinium eisenack & cookson 1960 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . trichodinium castanea deflandre 1935 ex clarke & verdier 1967 genus trithyrodinium drugg 1967 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . trithyrodinium? conservatum sp. nov. trithyrodinium evittii drugg 1967 trithyrodinium quinqueangulare marheinecke 1992 trithyrodinium suspectum (manum & cookson 1964) davey 1969b genus tuberculodinium wall 1967 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . tuberculodinium vancampoae (rossignol 1962) wall 1967 genus vesperopsis bint 1986 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . vesperopsis longicornis (batten & lister 1988) harding 1990 genus wallodinium loeblich jr. & loeblich iii 1968 . . . . . . . . . . . . . . . . . . . . . . . . . wallodinium luna (cookson & eisenack 1960a) lentin & williams 1973 genus wetzeliella eisenack 1938 emend. williams et al. in fensome et al. 2009 . . . . . wetzeliella articulata wetzel in eisenack 1938 emend. williams et al. in fensome et al. 2009 genus xenascus cookson & eisenack 1969 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . xenascus ceratioides (deflandre 1937) lentin & williams 1973 xenascus wetzelii slimani 1996 ex slimani 2001a systematics – acritarchs and other algae . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . genus fromea (cookson & eisenack 1958) yun hyesu 1981 . . . . . . . . . . . . . . . . . . . 9 67 68 69 69 70 71 71 72 73 73 73 74 75 75 76 76 76 77 77 bulletin36.qxp_bulletin 36 19/12/16 13.39 side 9 fromea nicosia jansonius 1989 fromea quadrangularis sp. nov. genus microsphaeridium benedek 1972 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . microsphaeridium ancistroides benedek 1972 genus palambages wetzel 1961 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . palambages spp. genus paralecaniella cookson & eisenack 1970a . . . . . . . . . . . . . . . . . . . . . . . . . . . . . paralecaniella indentata (deflandre & cookson 1955) cookson & eisenack 1970a genus pediastrum meyen 1829 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . pediastrum spp. genus tetraporina naumova 1950 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . tetraporina sp. a tetraporina sp. b systematics – miospores and fungal elements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . miospores . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . genus afropollis doyle et al. 1982 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . afropollis sp. genus appendicisporites weyland & krieger 1953 . . . . . . . . . . . . . . . . . . . . . . . . . . . . appendicisporites potomacensis brenner 1963 appendicisporites unicus (markova in ivanova & markova 1961) singh 1964 genus aquilapollenites rouse 1957 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . aquilapollenites quadrilobus rouse 1957 emend. braman 2013 genus azolla lamarck in lamarck et al. 1783 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . azolla spp. genus baculatisporites pflug & thomson in thomson & pflug 1953 . . . . . . . . . . . . . baculatisporites crenulatus sp. nov. genus callialasporites dev 1961 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . callialasporites dampieri (balme 1957) dev 1961 callialasporites obrutus norris 1969 genus caryapollenites raatz 1938 ex potonié 1960 . . . . . . . . . . . . . . . . . . . . . . . . . . . caryapollenites inelegans nichols & ott 1978 caryapollenites veripites (wilson & webster 1946) nichols & ott 1978 genus cerebropollenites nilsson 1958 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . cerebropollenites mesozoicus (couper 1958) nilsson 1958 genus chenopodipollis krutzsch 1966 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . chenopodipollis sp. genus cicatricosisporites potonié & gelletich 1933 . . . . . . . . . . . . . . . . . . . . . . . . . . cicatricosisporites minutaestriatus (bolkhovitina 1961) pocock 1964 cicatricosisporites ornatus srivastava 1972 genus cicatricososporites pflug & thomson in thomson & pflug 1953 . . . . . . . . . . . cicatricososporites eocenicus (selling 1944) jansonius & hills 1976 genus compositoipollenites potonié 1951 ex potonié 1960 . . . . . . . . . . . . . . . . . . . . . . compositoipollenites sp. b of williams & brideaux 1975 genus corsinipollenites nakoman 1965 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . corsinipollenites oculusnoctis (thiergart 1940) nakoman 1965 genus extratriporopollenites pflug in thomson & pflug 1952 ex pflug in thomson & pflug 1953 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . extratriporopollenites spp. genus graminidites cookson 1947 ex potonié 1960 . . . . . . . . . . . . . . . . . . . . . . . . . . graminidites sp. a. of williams & brideaux 1975 10 78 78 78 78 79 80 80 80 80 80 81 81 82 82 83 83 83 83 84 84 84 84 bulletin36.qxp_bulletin 36 19/12/16 13.39 side 10 genus momipites wodehouse 1933 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . momipites annellus nichols & ott 1978 momipites coryloides wodehouse 1933 genus osmudacidites couper 1953 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . osmundacidites wellmannii couper 1953 genus parviprojectus mtchedlishvili in samoilovitch & mtchedlishvili 1961emend. braman 2013 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . parviprojectus reticulatus mtchedlishvili in samoilovitch & mtchedlishvili 1961 genus parvisaccites couper 1958 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . parvisaccites amplus brenner 1963 parvisaccites radiatus couper 1958 genus periporopollenites pflug & thomson in thomson & pflug 1953 . . . . . . . . . . . . periporopollenites sp. genus pistillipollenites rouse 1962 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . pistillipollenites macgregorii rouse 1962 genus quercoidites potonié et al. 1950 ex potonié 1960 . . . . . . . . . . . . . . . . . . . . . . . quercoidites sp. genus rugubivesiculites pierce 1961 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . rugubivesiculites spp. genus tiliaepollenites potonié 1931 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . tiliaepollenites crassipites (wodehouse 1933) comb. nov. tiliaepollenites sp. a genus translucentipollis khlonova 1961 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . translucentipollis contiguus (tschudy 1969) braman 2013 genus wodehouseia stanley 1961 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . wodehouseia spinata stanley 1961 genus zlivisporis pacltová 1961 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . zlivisporis spp. genus zonalapollenites pflug in thomson & pflug 1953 . . . . . . . . . . . . . . . . . . . . . . zonalapollenites igniculus (potonié 1931) thomson & pflug 1953 fungal elements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . references . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . plates 1–20 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 11 86 84 85 85 85 86 86 86 86 87 87 87 87 88 88 89 103 bulletin36.qxp_bulletin 36 19/12/16 13.39 side 11 bulletin36.qxp_bulletin 36 19/12/16 13.39 side 12 13 abstract authors’ addresses r.a.f. & g.l.w., geological survey of canada (atlantic), natural resources canada, po box 1006, 1 challenger drive, dartmouth, nova scotia b2y 4a2, canada. e-mail: rob.fensome@canada.ca h.n.-h., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. new palynological analysis of samples from 13 offshore wells on the canadian margin and six wells on the west greenland margin has led to a new event biostratigraphic framework for cretaceous– cenozoic strata of the labrador sea – davis strait – baffin bay (labrador–baffin seaway) region. this framework is based on about 150 dinoflagellate cyst taxa and 30 acritarch, algal, fungal and plant microfossil (mostly miospore) taxa. in the systematics we include three new genera of dinocysts (scalenodinium, simplicidinium and taurodinium), 16 new species of dinocysts (chirop teridium gilbertii, chytroeisphaeridia hadra, cleistosphaeridium elegantulum, cleistosphaeridium palmatum, dapsilidinium pseudoinsertum, de flan drea borealis, evittosphaerula? foraminosa, gingi nodinium? flexidentatum, hystricho sphae ridium quadratum, hystrichostrogylon digitus, impleto sphaer idium apodastum, scalenodinium scalenum, surculo sphaeridium convocatum, talladinium pellis, taurodinium granulatum and trithyrodinium? conservatum), four emendations of dinocyst genera (alter bidinium, cha tangiella, chiropteridium and surculosphae ridium), six new combinations for dinocyst species (alterbidinium biaperturum, de flandrea majae, kleithriasphaeridium mantellii, simplicidinium insolitum, spongodinium grossum, spon godinium obscurum), one new acritarch species (fromea quadrangularis), one new miospore species (bacula ti s porites crenulatus) and one new combination for miospores (tiliaepollenites crassipites). most of the taxa included provide age information, almost exclusively last occurrences (range ‘tops’), but some are useful mainly for environmental interpretations. collectively, they provide a powerful tool for helping to establish the geological history of the labrador–baffin seaway. fensome, r.a., nøhr-hansen, h., & williams. g.l. 2016: cretaceous and cenozoic dinoflagellate cysts and other palynomorphs from the western and eastern margins of the labrador–baffin seaway. geological survey of denmark and greenland bulletin 36, 143 pp. bulletin36.qxp_bulletin 36 19/12/16 13.39 side 13 14 r canadacanada ellesmere island 70°n 60°n 45°w55°w65°w labrado r nuussuaq nuussuaq basin hareøen disko bylot island cape dyer scott inlet buchan gulf quqaluit/ padloping islands saglek b asin 646 112 113 645 umiivik-1 delta-1 t4-1 t8-1 alpha-1/s1 gamma-1 hellefisk-1 at7-1 at2-1 lf7-1 ikermiut-1 nukik-2 nukik-1 qulleq-1 gjoa g-37 ralegh n-18 hekja o-71 rut h-11 gilbert f-53 karlsefni a-13 skolp e-07 pothurst p-19 ogmund e-72 snorri j-90 bjarni o-82 bjarni h-81 herjolf m-92 north bjarni f-06 roberval k-92 north leif i-05 cartier d-70 hare bay e-21 freydis b-87 south labrador n-79 kangâmiut-1 647 hopedalebasin cum berland sound h om e bay b af f in is lan d canada d a v is s t ra i t la b ra d o r s e a gro#3 250 km baffin bay greenland h udson stra i t 75°w shallow cored borehole greenland–canada border odp or dsdp borehole exploration well with gas shows exploration well exploration well with oil and gas shows 500 km greenland labrador sea baffin bay d avis strait nar es st ra it b a ffin isla n d 60°n 60°w 40°w 70°n 80°n bulletin36.qxp_bulletin 36 19/12/16 13.39 side 14 15 canada and greenland are separated, from south to north, by the labrador sea, the davis strait, baffin bay (which we call collectively the labrador–baffin seaway), and the narrow nares strait (fig. 1). the labrador– baffin seaway stretches from about 52°n on to 75°n, a distance of roughly 2500 km. timing for the geological evolution of the seaway’s margin is based primarily on biostratigraphic analyses from exploration wells drilled between 1971 and 2000 on the labrador margin, off west greenland and in the davis strait, and from some shallow cored boreholes drilled in baffin bay in the 1980s (figs 1, 2). these drilling activities revealed thick succes sions of mesozoic–cenozoic sediments on both sides of the seaway, and the lithostratigraphy of these has developed in tandem with the biostratigraphy. additional in formation comes from numerous creta ce ous–cenozoic outcrop sections in the nuussuaq basin in west green land, a few onshore sections in labrador and on baffin and bylot islands (figs 1, 2), and from odp leg 105, hole 645 in baffin bay (fig. 1). although previous work on these materials has provided a good stratigraphic foundation for the cretaceous– cenozoic of the region (fig. 3), renewed petroleum exploration interest in recent years has revealed gaps in our knowledge and a need to correlate the western and eastern margins of the seaway. consequently, for the past decade, the geo logical survey of denmark and greenland (geus) and the geological survey of cana da (atlantic) (gsca) have undertaken an exhaustive paly nological study, based on palynological analysis of more than 2000 well samples. the objectives of this work have been to refine age control and palaeoenvironmental interpretations on a regional scale, with the ultimate goal of contribu ting fundamentally to an under standing of the geological history of the seaway and its petroleum systems. a significant finding has been that the orderly story of five regional unconformities (fig. 3) is not as clear-cut as has been previously thought (see fig. 4 and nøhr-hansen et al. 2016 for details). dinoflagellate cysts (dinocysts) are the primary palynomorph group evaluated, but spores and pollen (mio spores), a fern microspore massula (azolla), algal micro fossils and acritarchs have also been used. our new age determinations (figs 5, 6, in pocket; see nøhr-hansen et al. 2016 for further details) are more precise than those of previous studies because of advances made in refining stratigraphic ranges, especially of dinocysts (e.g. williams et al. 2004). moreover, the number of species for which we have detailed stratigraphic information has increased immeasurably since early studies in the region in the 1970s. this paper focuses on the systematic treatment of palynomorphs observed in the study, thereby underpinning the results presented in the companion bulletin (nøhr-hansen et al. 2016), but also facilitating future work in the region and further afield. introduction facing page: fig. 1. map of the labrador–baffin seaway showing the location of relevant wells, boreholes and onshore localities; wells and boreholes shown in red were used in this study. the cored borehole ggu 400712 is 2 km ene of umiivik-1; annertuneq is on the north coast of nuussuaq. inset map shows the regional context of the labrador–baffin seaway between canada and greenland. bulletin36.qxp_bulletin 36 19/12/16 13.39 side 15 16 rre s st es s t ai t ai t ellesmere island ■■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ▲ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■ ■ ■ ■ ■■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■ ■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■■ ■■ ■■ ■ ■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■■ ■ ■■ ■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■■ 646 112 113 645 rut h-11 gilbert f-53 karlsefni a-13 pothurst p-19 skolp e-07 ogmund e-72 snorri j-90 bjarni o-82 roberval k-92 north leif i-05 freydis b-87 cartier d-70 herjolf m-92 north bjarni f-06 250 km south labrador n-79 647 saglek basin hellefisk-1 ikermiut-1 nukik-2 delta-1 t4-1 t8-1 gamma-1 alpha-1/s1 nukik-1 qulleq-1 lf7-1 at2-1 at7-1 gjoa f-37 ralegh n-18 hekja o-71 kangâmiut-1 nuussuaq svartenhuk halvø disko cape dyer umiivik-1 gro#3 45°w 35°w sediments, large basins (400–0 ma) oceanic crust, eocene oceanic crust, age uncertain basalts and intrusives (60–30 ma) basement locally covered by sediment spreading axis extensional fault compressional fault thrust transition fault odp or dsdp borehole oceanic crust, paleocene exploration well exploration well with gas shows, exploration well with oil and gas shows shallow cored borehole fault (undifferentiated) paam iut b asin n u u k b as in k an gâ m iu t b as in lady franklin basin sisimiut basin disko bugt ik er m iu t b as in h o m e b ay b asin lancaster basin eclipse trough m elville bay g raben buchan g raben scott g raben d av is s tr ai t h ig h baffin basin n u u ss u au q b as in labrado r labrador sea hudson strait c um berland sound m elville bay k ivioq basin canada greenland bylot island baffin island 65°w 55°w 45°w 60°n 70° 70°n 55°w65°w75°w h opedale basin hawke basin west greenland volcanic province bulletin36.qxp_bulletin 36 19/12/16 13.39 side 16 17 the study is based on samples from ditch cuttings, sidewall cores and conventional cores from onshore and offshore wells and boreholes, and on outcrop samples (figs 1, 2). most of the samples were processed for palynology between 1990 and 2006, using standard techniques for concentrating palynomorphs. however, many of the sidewall-core samples are represented only by oil-company slides produced in the late 1970s to 1980s and are of variable quality. using transmitted light microscopy, qualitative and quantitative analyses were carried out of the dinocysts, acritarchs and pollen and spores, the results and implications of which are detailed in nøhr-hansen et al. (2016). the systematic section below is accompanied by plates; sample number, slide number and england finder coordinates for each specimen are provided. the slides from the six offshore west greenland wells, hellefisk-1, ikermiut-1, kangâmiut-1, nukik-1, nukik2 and qulleq-1 are kept at geus. geus also processed samples from six offshore canadian wells, gjoa g-37, hekja o-71, north leif i-05, ogmund e-72, ralegh n-18 and skolp e-07. one set of palynology slides from gjoa g-37, hekja o-71 and ralegh n-18 is housed at geus (copenhagen) and one set is housed at the geological survey of canada (atlantic), dartmouth, nova scotia, canada. a set from each well is also curated at the canada–nova scotia offshore petroleum board, dartmouth, nova scotia. the geological survey of canada has processed samples from all the labrador margin wells, which include bjarni o-82, gilbert f-53, kalsefni a-13, roberval k-92, rut h-11, snorri j-90 and south labrador n-79. one set of palynology slides from these wells is housed at the geological survey of canada (atlantic), dartmouth, nova scotia. another set is curated at the canada – newfoundland and labrador offshore petroleum board (cnlopb) in st. johns, newfoundland. also on file with cnlopb are sets of geus-prepared slides from north leif i-05, ogmund e-72 and skolp e-07. facing page: fig. 2. map of the basic geology of the labrador–baffin seaway, showing the main structural features and relevant well/borehole locations; wells and boreholes shown in red were used in this study. modified from henriksen et al. (2009) and oakey & chalmers (2012). methodology bulletin36.qxp_bulletin 36 19/12/16 13.39 side 17 18 upper cretaceous paleocene eocene oligocene miocene pliocene quaternary lower cretaceous 0 ma 5 10 15 20 25 30 35 40 45 50 55 60 65 70 75 80 85 90 95 100 105 110 115 120 125 130 135 140 145 aptian barremian hauterivian valanginian berriasian albian cenomanian turonian coniacian santonian campanian maastrichtian danian selandian thanetian ypresian lutetian bartonian priabonian rupelian chattian aquitanian burdigalian langhian serravallian tortonian messinian zanclean piacenzian c28 c29 c30 c31 c32 c33 c34n c34n c34n c34n m"-3"r m"-2"r m"-1"r c27 c26 c25 c24 c22 c21 c20 c19 c18 c17 c23 c16 c15 c13 magnetostrat. (polarity, chron) chronostratigraphy unconformitiesshelf basin labrador marginsw ne labrador trough erosion leif mb upper freydis mb lower freydis mb upper and middle gudrid mb lower gudrid mb sa gle k f m mokami fm lower kenamu fm cartwright fm lower lower upper upper upper bjarni fm alexis fm markland fm (labrador u nc.) (avalon unc.) (bylot unc.) bylot avalon labrador baffin bay beaufort (baffin bay unc.) (b eaufort u n c. ) fig. 3. stratigraphic framework of the mesozoic–cenozoic rocks of the labrador and west green land margins and adjacent on shore sections; modified from gregersen et al. (2013). the la brador margin stratigraphy is from dickie et al. (2011), with unconformities from mcwhae (1981). the west greenland mar gin stratigraphy is based on well stratigraphic studies by rolle (1985), nøhr-hansen (2003) and sønderholm et al. (2003); the deeper sub-well section is based on chalmers et al. (1993), chal mers & pulvertaft (2001) and sørensen (2006). the nuussuaq basin stratigraphy is from storey et al. (1998), dam et al. (2009), pedersen & nøhr-hansen (2014) and larsen et al. (2015); these pa pers and this study are the source of the inferred ages. the southeast baffin island stratigraphy is from burden & langille (1990) and pedersen et al. (2002), the home bay and scott inlet seabed samples from maclean et al. (2014) and the north baffin is land stratigraphy is based on jack son et al. (1978) and mcwhae (1981). the bylot island strati graphy is from miall (1986), waterfield (1989), and harrison et al. (1999). the timescale (ma) and magnetostratigraphy are from gradstein et al. (2012). fm: for mation (formal). fm: formation (informal). mb: member (for mal). mb: member (informal). unc.: unconformity. bulletin36.qxp_bulletin 36 19/12/16 13.39 side 18 19 h el le fis k1 g r o # 3 u m iiv ik -1 ik er m iu t1 k an gâ m iu t1 n u ki k2 n u ki k1 q u lle q -1 west greenland margin kangilia fm nuussuaq basin se baffin islandnwses n bylot island itilli fm aaffarsuaq mb atane fm ? (ikermiut fm) kangeq sequence ikermiut fm ikermiut fm narssarmiut fm kangâmiut fm kangâmiut fm kangâmiut fm kangâmiut fm hellefisk fmnukik fm nukik fm ? ? ? ? ? ? ? ? ? ? ? ataneq fm ataneq fm manîtsoq fm manîtsoq fm manîtsoq fm fylla sand erqua fm naqerlog fm hareøen fm (hareøen) youngest volcanics quqaluit fm cape searle fm n u u ss u aq g ro u p (home bay – scott inlet) cape dyer basalt maligât fm svartenhuk fm vaigat fm atanikerluk fm eqalulik fm quikavsak fm kome fm basalts? sills/dykes appat sequence kitsissut sequence ? ? upernivik næs fm slibestensfjeldet fm hassel fm (labrador unc.) sermilik fm kanguk fm aktineq fm pond inlet fm navy board fm bylot island fm agatdal fm ikermiut fm mainly marine mudstones, locally sandy/silty mainly marine or deltaic sandy/silty deposits, locally mudstones potential and known source rocks volcanics td in basement unconformity mainly continental deposits bulletin36.qxp_bulletin 36 19/12/16 13.39 side 19 20 0 ma 5 10 15 20 25 30 35 40 45 50 55 60 65 70 75 80 85 90 95 100 105 110 115 120 125 late cretaceous paleocene eocene oligocene miocene pliocene quaternary p er io d / e p o ch a ge ( st ag e) fo rm at io n m em b er r eg io n al (d ic ki e et a l. 2 0 1 1 ) l ab ra d o r m ar gi n (m cw h ae 1 9 8 1 ) early cretaceous aptian albian cenomanian turonian coniacian santonian campanian maastrichtian danian selandian thanetian ypresian lutetian bartonian priabonian rupelian chattian aquitanian burdigalian langhian serravallian tortonian messinian zanclean piacenzian mokami kenamu cartwright markland bjarni saglek upper lower upper leif lower lower upper + middle gudrid lower gudrid upper freydis lower freydis upper sn o rr i j9 0 h e k ja o -7 1 n o rt h l e if i0 5 b e au fo rt b af fin b ay b yl o t a va lo n l ab ra d o r r o b e rv al k -9 2 b ja rn i o -8 2 o gm u n d e -7 2 s. l ab ra d o r n -7 9 sk o lp e -0 7 k ar ls e fn i a -1 3 g ilb e rt f -5 3 r u t h -1 1 r al e gh n -1 8 g jo a g -3 7 s n chronostratigraphy hiatuses labrador margin preserved stratigraphic interval labrador margin lithostratigraphy bulletin36.qxp_bulletin 36 19/12/16 13.39 side 20 21 ma h e lle fis k -1 n u k ik -1 n u k ik -2 k an gâ m iu t1 ik e rm iu t1 q u lle q -1 g r o # 3 & s u rf ac e se ct io n s n . n u u ss u aq u m iiv ik -1 , sv ar te n h u k h al vø & n . n u u ss u aq , su rf ac e s e ct io n s s n late cretaceous paleocene eocene oligocene miocene pliocene quaternary p er io d / e p o ch a ge ( st ag e) early cretaceous aptian albian cenomanian turonian coniacian santonian campanian maastrichtian danian selandian thanetian ypresian lutetian bartonian priabonian rupelian chattian aquitanian burdigalian langhian serravallian tortonian messinian zanclean 0 5 10 15 20 25 30 35 40 45 50 55 60 65 70 75 80 85 90 95 100 105 110 115 120 125 n o t p en et ra te d o ff sh o re chronostratigraphy west greenland margin preserved stratigraphic interval stratigraphic record labrador 0 % % 100 100 west greenland (offshore) 0 piacenzian fig. 4. ages of strata preserved in wells on the labrador, baffin and green land margins, compared with the hiatuses linked to uncon formities reported in the literature. dickie et al. (2011) recognised a number of unconformities in two or more of three regions – the southwest greenland shelf, the labrador shelf and the jeanne d’arc basin, offshore newfound land (red, all three regions; green, two regions). mcwhae (1981) named five unconformities on the labrador margin; the hiatuses associated with these unconformities indicated here are estimates by mcwhae (1981) based on an unspecified timescale and thus should be considered approximate. the ‘cumulative preservation’ plots (right) depict sche ma tically the proportion of wells that contain sediments of a specified age, based on one millionyear slots; note that these cal culations only include wells that extend down to/beyond the relevant stratigraphic level. bulletin36.qxp_bulletin 36 19/12/16 13.39 side 21 22 phelodinium apectodinium,axiodinium, charlesdowniea, petalodinium, piladinium, rhombodinium, sophismatia, stichodinium, talladinium, wetzeliella alterbidinium, cerodinium, chatangiella, deflandrea, isabelidinium, lentinia, palaeocystodinium, scalenodinium?, senegalinium, spinidinium, trithyrodinium ginginodinium?, laciniadinium, palaeoperidinium, phthanoperidinium? subtilisphaera atopodinium, batiacasphaera, chlamydophorella, impletosphaeridium?, raphidodinium, reticulatosphaera, simplicidinium, tanyosphaeridium, wallodinium alisocysta, eocladopyxis, homotryblium hystrichosphaeridium heteraulacacysta tuberculodinium aptea, nyktericysta, odontochitina, pseudoceratium, taurodinium, vesperopsis, xenascus adnatosphaeridium, areoligera, areosphaeridium, chiropteridium, cleistosphaeridium, enneadocysta, glaphyrocysta, heterosphaeridium, licracysta palynodinium, schematophora, senonaisphaera, tenua callaiosphaeridium, cerebrocysta, chytroeisphaeridia, fibrocysta, habibacysta, kiokansium, surculosphaeridium, trichodinium cannosphaeropsis, evittosphaerula, hystrichosphaeropsis, hystrichostrogylon, impagidinium, rottnestia, spiniferites, achilleodinium, apteodinium, cordosphaeridium, cribroperidinium, dapsilidinium, diphyes, disphaerogena, hapsocysta, hystrichokolpoma, lingulodinium, operculodinium, spongodinium, thalassiphora eatonicysta, kleithriasphaeridium, oligosphaeridium gillinia, histiocysta batioladinium uncertain cribroperidinioideae gonyaulacoideae uncertain helgolandinioideae goniodomoideae protoperidinioideae podolampaceae peridiniaceae ceratiaceae ceratocoryaceae areoligeraceae heterocapsaceae heterodiniaceae uncertain crypthecodiniaceae pyrocystaceae goniodomaceae pyrodinioideae gambierdiscoideae glenodiniaceae palaeoperidinioideae deflandreoideae calciodinelloideae ovoidinioideae lithoperidinioideae peridinioideae wetzelielloideae leptodinioideae diplopsalioideae protoperidinioideae gonyaulacaceae rhaetogonyaulacaceae cladopyxiaceae pareodiniaceae pareodinioideae broomeoideae o rd er g o n ya u la ca le s o rd er p er id in al es peridiniineae ceratiineae gonyaulacineae uncertain goniodomineae heterocapsineae glenodiniineae cladopyxiineae rhaetogonyaulacineae family subfamily generasuborder fig. 7. suprageneric affiliation of peridiniphycidean genera treated systematically in the present study. the only non-peridiniphycidean dinocyst genus treated in this study is dinogymnium, which belongs to the subclass gymnodiniophycidae, order ptychodiscales, family ptychodiscaceae, subfamily dinogymnioideae. bulletin36.qxp_bulletin 36 19/12/16 13.39 side 22 23 division dinoflagellata (bütschli 1885) fensome et al. 1993 class dinophyceae pascher 1914 genus achilleodinium eaton 1976 type. eisenack 1954, plate 11, fig. 18, as hystricho sphaer idium biformoides. 1976 achilleodinium eaton, p. 234. remarks. we follow the synopsis of fensome et al. (2009, p. 11) for this genus. achilleodinium biformoides (eisenack 1954) eaton 1976 (plate 1, figs 1–3) 1954 hystrichosphaeridium biformoides eisenack, p. 68, plate 11, figs 16–20. 1965 baltisphaeridium biformoides (eisenack) – downie & sarjeant, p. 87. 1965 hystrichokolpoma biformoides (eisenack) – rozen, p. 308. 1976 achilleodinium biformoides (eisenack) – eaton, p. 234. 1980 florentinia biformoides (eisenack) – duxbury, p. 121. age. lo: ypresian. genus adnatosphaeridium williams & downie 1966a type. williams & downie 1966a, plate 24, fig. 7, textfig. 56, as adnatosphaeridium vittatum. 1966a adnatosphaeridium williams & downie, p. 215. remarks. stancliffe & sarjeant (1990, p. 199–200) emen ded the diagnosis of adnatosphaeridium but did not change the concept of the generic circumscription. here, we follow the synopsis provided by fensome et al. (2009, p. 11). systematics – general systematics – dinoflagellates presentation of the systematic taxonomic descriptions includes subsections on dinocysts, acritarchs and other algae, miospores, and fungi. for pragmatic reasons, genera are arranged alphabetically within each section. for the dinocysts, a tabulation of suprageneric affinities is provided (fig. 7). under individual taxa, all synonyms are listed but those already indicated and credited to earlier works in fensome et al. (2008) are generally not discussed further. the term ‘synopsis’ is used for a statement of what we consider the essential defining features of the genus. unless otherwise clearly indicated, the synopses given herein should not be understood to represent emendations. only those emendations that we agree with or consider helpful are cited. all ages refer to their definition in the latest revision of the geological timescale (gradstein et al. 2012). for most species, an age is indicated for significant events in its stratigraphical range, usually a range top or ‘last occurrence’ (lo). the statement ‘not plotted’ refers to the fact that a particular taxon was not plotted in the summary event charts (figs 5, 6, in pocket); the reasons for not plotting some taxa include a taxon’s sparsity or long and unhelpful range (though it may be helpful palaeoenvironmentally. because the data are based main ly on cuttings samples, a taxon’s range bottom or ‘first occurrence’ (fo) is rarely used. we occasionally refer to the lad of a species, its ‘last appearance datum’, which refers to its youngest global occurrence. bulletin36.qxp_bulletin 36 19/12/16 13.39 side 23 24 adnatosphaeridium vittatum williams & downie 1966a (plate 1, figs 4, 8) 1966a adnatosphaeridium vittatum williams & downie, p. 15, plate 24, fig. 7; text-fig. 56. 1966a adnatosphaeridium multispinosum williams & downie, p. 216, 217, plate 24, fig. 5; text-fig. 57. age. lo: late bartonian. not plotted. remarks. following fensome et al. (2009, p. 13), the species adnatosphaeridium multispinosum is considered a taxonomic synonym of adnatosphaeridium vittatum. genus alisocysta stover & evitt 1978 type. drugg 1967, plate 1, fig. 12, as eisenackia cir cum tabulata. 1978 alisocysta stover & evitt, p. 15, 16. 1979a agerasphaera harland, p. 28, 29; illegitimate name, having the same type as alisocysta. synopsis. goniodomacean (pyrodinioid) cysts that are proximate and subspherical. the tabulation is reflected by penitabular ridges or septa that delineate the plates, including the cingulars and some or all of the sulcals. the archaeopyle is apical, with the formula a(1–4´); operculum free. remarks. quattrocchio & sarjeant (2003, p. 144) considered alisocysta to be a taxonomic junior synonym of eisenackia. however, as stated by stover & evitt (1978, p. 42) “alisocysta has parasutural [i.e. penitabular] ridges or septa rather than depressions as in eisenackia” and alisocysta is thus retained here as a separate genus, including both the type, alisocysta circumtabulata, and alisocysta margarita. alisocysta circumtabulata (drugg 1967) stover & evitt 1978 (plate 1, figs 6, 7) 1967 eisenackia circumtabulata drugg, p. 15, plate 1, figs 12, 13. 1978 hystrichokolpoma circumtabulatum (drugg) – schumacker-lambry, p. 42. 1978 alisocysta circumtabulata (drugg) – stover & evitt, p. 16. 1979a agerasphaera circumtabulata (drugg) – harland, p. 29; illegitimate combination as the generic name agerasphaera is illegitimate. 2003 eisenackia circumtabulata (drugg) – quattrocchio & sarjeant, p. 146. age. lo: earliest thanetian. alisocysta margarita (harland 1979a) harland 1979a (plate 1, figs 5, 9) 1979a agerasphaera margarita harland, p. 29, 31, 33, plate 1, figs 1–12; plate 2, figs 1–10. 1979a alisocysta margarita (harland) – harland, p. 35. 2003 eisenackia margarita (harland) – quattrocchio & sarjeant, p. 146. age. lo: earliest thanetian. genus alterbidinium lentin & williams 1985 emend. nov. type. vozzhennikova 1967, plate 77, fig. 2, as albertia recticornis. 1967 albertia vozzhennikova, p. 150, 151; illegitimate name. 1976 alterbia lentin & williams, p. 47, 48; illegitimate name. 1985 alterbidinium lentin & williams, p. 14. emended description. peridiniacean (deflandreoid) cysts that are proximate and peridinioid, usually elongate, in outline. the antapical horns are always asymmetrically arranged, the left horn being larger. bicavate. the pericyst surface is generally atabulate, smooth, or with low ornament; the cingulum is commonly indicated, if only marginally. the periarchaeopyle is intercalary or combination intercalary–precingular; always involving an isoto stenodeltaform hexa plate 2a and commonly plate 4´´, the operculum remaining attached posteriorly; archaeopyle i2a @ or (i2ap4´´)@. remarks. in their emendation of alterbidinium, khowajaateequzzaman et al. (1991, p. 38) stated: “archaeopyle intercalary, independently developed on periphragm and endophragm, dissimilar in shape; periarchaeopyle hexa 2a, steno/iso-deltaform, perioperculum free or adnate; endoarchaeopyle hexa 2a, eury-deltaform, endoperculum adnate.” however, they did not base their emendation on bulletin36.qxp_bulletin 36 19/12/16 13.39 side 24 the type of the genus. thus, in the emendation proposed here, the synop sis of fensome et al. (2009, p. 13) is largely repeated: it emphasised the nature of the archaeopyle, which is formed from the isoto stenodeltaform hexa plate 2a, but may also involve plate 4´´. almost invariably, the operculum remains attached posteriorly. genera with similar morphologies to alterbidinium include spinidinium, diconodinium, cerodinium, cha tan giella and isabelidinium. spinidinium and dicono dinium differ from alterbidinium in having a spinate peri ph ragm. cerodinium has a large isodeltaform periarchaeopyle, an endoarchaeopyle formed from the loss of one to three anterior intercalary plates, and generally long apical and antapical horns. chatangiella is distinguished by its gen erally omegaform 2a and partite cingulum, although fen some et al. (2009, p. 19) stated that the archaeopyle could vary between isodeltaform and iso-omegaform. we now consider that chatan giella should be restricted to taxa that possess an isoor lati-omegaform archaeopyle. sim ilarly, isabelidinium is restricted to taxa with an isoto latiomega form 2a plate and archaeopyle. however, some species presently included in isabelidinium can have the taform or isodeltaform 2a plates and archaeopyles. the present morphological differences used to separate the deflandreoid genera are in large part unsatisfactory. perhaps a more effective approach would be to adopt the methodological approach used by williams et al. (2015) in their reclassification of the wetzelielloideans. following this example, the deflandroidean genera would be differentiated primarily on the nature of the archaeopyle, since this reflects differences in tabulation pattern. this then follows the categories recognised by lentin & williams (1976) and expanded by bujak & davies (1983). these are omega form, thetaform and deltaform modes, with steno-, isoand latisubcategories. such an approach would not be a perfect solution as there are clearly gradations, but it would make the generic distinctions more meaningful, and is considered here the best solution. in the spirit of this approach, herein we informally use isabelidinium for forms having a latito iso-omegaform archaeopyle. a formal revision would require many of the species now included in that genus to be transferred to other existing or new genera. indeed, a re-appraisal of all deflandreoidean taxa is deemed necessary, but is beyond the scope of the present paper, although as an initial step emendations are proposed for alterbidinium and chatangiella. alterbidinium acutulum (wilson 1967a) lentin & williams 1985 (plate 1, figs 10, 11) 1967a deflandrea acutula wilson, p. 225, 226, figs 11, 12. 1967 albertia curvicornis vozzhennikova, p. 151, plate 76, figs 1–4. 1967 albertia recticornis vozzhennikova, p. 151, 152, plate 77, figs 1–4; plate 78, figs 1–3; plate 79, figs 1, 2. 1976 alterbia acutula (wilson) – lentin & williams, p. 48. 1976 alterbia curvicornis (vozzhennikova) – lentin & williams, p. 49. 1976 alterbia recticornis (vozzhennikova) – lentin & williams, p. 47. 1985 alterbidinium acutulum (wilson) – lentin & williams, p. 14. age. lo: early maastrichtian. remarks. fensome & williams (2005, p. 9) recorded this species, in part, under the informal name ‘alterbidinium fleximorphum’ because of the great variation shown by this species. it is now recognised that this variability should be encompassed within alterbidinium acutulum. alterbidinium biaperturum (mcintyre 1975) comb. nov. (plate 1, figs 12–14) 1975 deflandrea biapertura mcintyre, p. 66, plate 3, figs 5–8. 1976 chatangiella? biapertura (mcintyre) – lentin & williams, p. 53. age. lo: late maastrichtian. remarks. this species is distinguished by an opening in the pericyst between the antapical horns. from the illustrations of the holotype in mcintyre (1975, plate 3, figs 5, 6), alterbidinium (as chatangiella?) biaperturum ap pears to possess a cingulum delineated by continuous ridges; the periarchaeopyle is extremely unusual, being distinctly deltaform on one lateral margin but, as m. pearce (personal communication 2015) has noted, possibly thetaform on the other lateral margin. whether such asymmetry occurs in other specimens of this species is unknown. but regardless, the species is thus not assign able to chatangiella. lentin & williams (1976) were aware of these differences when provisionally including this taxon in chatangiella, but did not consider assigning it to alterbidinium. we herein transfer the species to alter bidinium to more accurately reflect its morphology. 25 bulletin36.qxp_bulletin 36 19/12/16 13.39 side 25 26 alterbidinium? bicellulum (islam 1983a) lentin & williams 1985 (plate 1, figs 15, 16) 1983a alterbia? bicellula islam, p. 335, 336, plate 1, figs 6, 7. 1985 alterbidinium? bicellulum (islam) – lentin & williams, p. 14. age. lo: lutetian, with a peak occurrence in the ypresian. remarks. in the diagnosis for alterbidinium? bicellulum (as alterbia? bicellula), islam (1983a, p. 336) stated that the archaeopyle was “… intercalary type i/i with standard hexa style operculum attached”. this implies that the archaeopyle is isodeltaform, but it is not possible to confirm this from his illustrations (islam 1983a, plate 1, figs 6, 7, 10, 11). islam (1983a) questionably assigned this species to alterbia (now alterbidinium) “because the epipericoel is not always communicative to the exterior, and its degree of cavation and the archaeopyle index do not match those prescribed for the genus. the degree of cavation and the archaeopyle index of this species also do not match those of other peridinioid genera that are differentiated on the basis of these features by lentin & williams (1976).” taxa with an isodeltaform archaeo pyle and an attached operculum are rare, especially in the cenozoic. thus we leave the species questionably in alberbidinium. alterbidinium ioannidesii pearce 2010 (plate 1, fig. 17) 1986 dinoflagellate type e of ioannides, p. 42, plate 23, figs 13–16. 2010 alterbidinium ioannidesii pearce, p. 66, 67, plate 1, figs 1–6. age. lo: early campanian. remarks: this is an unusual species of alterbidinium in displaying tabulation on the pericyst. alterbidinium varium kirsch 1991 (plate 1, figs 18, 19) 1991 alterbidinium varium kirsch, p. 98, 99, plate 19, figs 1–10; text-figs 46a–h, 47a, b. age. lo: campanian. genus apectodinium (costa & downie 1976) lentin & williams 1977a emend. williams, damassa, fensome & guerstein in fensome et al. 2009 type. deflandre & cookson 1955, plate 5, fig. 7, as wetzeliella homomorpha. 1976 wetzeliella subgenus apectodinium costa & downie, p. 608. 1977a apectodinium (costa & downie 1976) – lentin & williams, p. 8. 2009 apectodinium (costa & downie) – emend. williams, damassa, fensome & guerstein in fensome et al., p. 13, 14. remarks. the emended diagnosis in fensome et al. (2009, p. 13, 14) is followed here; this emphasised archaeopyle type and variability and the general absence of a pericoel except in the vicinity of the horns. apectodinium homomorphum (deflandre & cookson 1955) lentin & williams 1977a (plate 1, fig. 20) 1955 wetzeliella homomorpha deflandre & cookson, p. 254, plate 5, fig. 7; text-fig. 19. 1974 hystrichosphaeridium caiobense regali et al., p. 290, plate 24, fig. 4. 1977a apectodinium homomorphum (deflandre & cookson) – lentin & williams, p. 8. 1981 apectodinium caiobense (regali et al.) – lentin & williams, p. 14. 1983a apectodinium folliculum islam, p. 336, 337, plate 1, figs 8, 9. age. peak: early ypresian. remarks. harland (1979b, p. 64) emended the diagnosis of this species, as wetzeliella (apectodinium) homomorphum, to “take into account the nature of the archaeo pyle, cingular and sulcal details”. a characteristic feature of apectodinium homomorphum is the absence of horns, although harland noted that incipient horns may be developed in the apical, lateral and antapical areas. apectodinium parvum (alberti 1961) lentin & williams 1977a (plate 2, figs 1–3) 1961 wetzeliella parva alberti, p. 8, 9, plate 1, figs 14–18; plate 12, figs 10–12. bulletin36.qxp_bulletin 36 19/12/16 13.39 side 26 27 1976 wetzeliella subgenus apectodinium parva (alberti) – costa & downie, p. 608. 1977a apectodinium parvum (alberti) – lentin & williams, p. 9. age. lo: earliest ypresian. remarks. in his emendation of apectodinium parvum, harland (1979b, p. 65, 66) noted that lateral horns are absent and the processes may be intratabular or sutural. in his description he stated that “the horn development, prominent at the antapex, is variously developed with relation to the apical horn, which may or may not be present.” the labrador margin specimens usually have an apical horn, but show considerable variation in length and width. some of the labrador margin specimens have longer processes than those illustrated by harland (1979b), with those of one specimen being 21 μm. when longer, the processes are extremely delicate. however, forms with more robust processes also occur. apectodinium quinquelatum (williams & downie 1966b) costa & downie 1979 (plate 2, fig. 4) 1948 hystrichosphaeridium geometricum pastiels, p. 41, plate 4, figs 1–5, 7–10, non hystricho sphaeridium geometricum deflandre 1942. 1966b wetzeliella homomorphum var. quinquelata williams & downie, p. 191, 192, plate 18, fig. 7. 1977a wetzeliella homomorpha subsp. quinquelata (williams & downie) – lentin & williams, p. 8. 1979b wetzeliella (apectodinium) quinquelata (williams & downie) – harland, p. 67. 1979 apectodinium quinquelatum (williams & downie) – costa & downie, p. 43. age. lo: ypresian. not plotted. genus aptea eisenack 1958 type. eisenack 1958, plate 22, fig. 5, as aptea polymorpha. 1958 aptea eisenack, p. 393. 1966a doidyx sarjeant, p. 205. aptea polymorpha eisenack 1958 (plate 2, fig. 5) 1958 aptea polymorpha eisenack, p. 394, plate 22, figs 5–12; plate 24, fig. 5. 1986 pseudoceratium polymorphum (eisenack 1958) – bint, p.145. age. lo: aptian. remarks. the ornamentation of this species is variable, ranging from coarsely reticulate to irregularly spinate. genus apteodinium eisenack 1958 type. eisenack 1958, plate 23, fig. 9, as apteodinium granulatum. 1958 apteodinium eisenack, p. 385. 1961 emslandia gerlach, p. 171. 1971 coniferatium burgess, p. 80, 81. remarks. the synopsis for apteodinium followed here is that provided in fensome et al. (2009, p. 14), which noted that tabulation may sometimes be weakly reflected on the cyst wall. apteodinium australiense (deflandre & cookson 1955) williams 1978 (plate 2, fig. 6) 1955 gymnodinium australiense deflandre & cookson, p. 248, plate 5, fig. 1. 1961 gonyaulax tenuitabulata gerlach, p. 159, plate 25, figs 10, 11; text-figs 1–3. 1965 emslandia australiensis (deflandre & cookson) – nagy, p. 202; combination not validly published. 1965a scriniodinium australiense (deflandre & cookson) – cookson & eisenack, p. 122. 1969 gonyaulacysta tenuitabulata (gerlach) – de coninck, p. 23. 1978 apteodinium australiense (deflandre & cookson) – williams, p. 794. 1978 millioudodinium tenuitabulatum (gerlach) – stover & evitt, p. 174. 1981 emslandia crassimurata benedek & sarjeant, p. 320, 322, fig. 1, nos 2, 4. 1984 cribroperidinium tenuitabulatum (gerlach) – helenes, p. 124. 1984a rhynchodiniopsis tenuitabulata (gerlach) – sarjeant, p. 76. age. lo: rupelian. bulletin36.qxp_bulletin 36 19/12/16 13.39 side 27 28 remarks. apteodinium australiense has a spongy wall, as does the holotype of cribroperidinium tenuitabulatum. moreover, the latter has at best weakly developed and barely discernible parasutural ridges. hence, we consider cribroperidinium tenuitabulatum to be a junior synonym of apteodinium australiense. apteodinium spiridoides benedek 1972 (plate 2, figs 7, 8) 1972 apteodinium spiridoides benedek, p. 5, plate 2, fig. 1a, b; plate 15, figs 1–6. 1981 emslandia spiridoides (benedek) – benedek & sarjeant, p. 318. age. lo: earliest serravallian. remarks. in their emendation of apteodinium spiridoides, benedek & sarjeant (1981, p. 318, 319) noted the presence of an apical horn, the two-layered nature of the wall and the archaeopyle type. jan du chêne et al. (1986, p. 48) retained the species in apteodinium. genus areoligera lejeune-carpentier 1938 type. lejeune-carpentier 1938, text-fig. 2, as areoli gera senonensis. 1938 areoligera lejeune-carpentier, p. b164. remarks. we agree with the synopsis for areoligera pro vided by fensome et al. (2009, p. 14) and their observations under remarks. included in areoligera are several species with overlapping morphology. this is primarily manifested in the shape and distribution of the process complexes and the complexity of the linkages between processes within an individual complex. fensome et al. (2009, p. 15) introduced two terms for the distribution of the process complexes. the encircling condition is when complexes are, “…arcuate to annulate and occur on all the preand postcingular plates.” in the disjunct condition, “…annulate to arcuate complexes are restricted to the dorsal surface, there is lateral development of linear complexes and the ventral surface (i.e. 6´´ and 6´´´) is devoid of complexes.” fensome et al. (2009, p. 15) define areoligera gippingensis as a species with “encircling complex arrangement and predominantly trabecular connections.” following the specific differences re cognised for species of areoligera by these authors, it is concluded that all specimens with this general morphology studied here should be assigned to areoligera gippingensis. the genera chiropteridium and glaphyrocysta superficially resemble areoligera, but neither chiropteridium nor glapyrocysta have distinct annulate or arcuate process complexes with basal ridges that delineate tabulation. however, chiropteridium gilbertii does have intratabular processes or complexes. chiropteridium differs from areo ligera primarily in having neither mid-dorsal nor midventral processes or membranes. glaphyrocysta has non tabular or indistinctly contabular processes. areoligera circumsenonensis fensome et al. 2009 (plate 2, fig. 9) 1966a areoligera cf. senonensis of williams & downie, p. 230, 231; text-fig. 64 a–c (not plate 25, fig 6, typographically mislabelled plate 26, fig. 6). 1969 areoligera senonensis lejeune-carpentier – gocht, p. 56, plate 8, fig. 4a, b (?not figs 5– 9); text-figs 40a, b, ?40c, d (not fig. 40e, f). 2009 areoligera circumsenonensis fensome et al., p. 15, plate 1, fig. m. age. lo: early lutetian. not plotted. remarks. areoligera circumsenonensis has processes like those of areoligera senonensis but the process complexes are annular to arcuate on the large preand postcingular plates. in areoligera gippingensis, the processes are interconnected along their length within, and occasionally between, complexes. areoligera gippingensis jolley 1992 (plate 2, figs 10, 11) 1966a areoligera cf. medusettiformis of williams & downie, p. 229, plate 25, fig. 4. 1976 areoligera cf. medusettiformis of eaton, p. 246, plate 3, fig. 7. 1992 areoligera gippingensis jolley, p. 26, 28, 30, 31, plate 1, figs 1–6; plate 2, figs 1–6; textfigs 2a–d, 3. age. lo: late ypresian. peak: early thanetian. remarks. specimens of areoligera gippingensis in labrador margin samples show considerable variability in the delineation of the process complexes. we have found that this species is gradational with glaphyrocysta divaricata. the species can be common in the late thanetian, but generally peaks in the early thanetian. bulletin36.qxp_bulletin 36 19/12/16 13.39 side 28 29 genus areosphaeridium eaton 1971 type. klumpp 1953, plate 18, figs 3, 4, as hystricho sphaeridium dictyoplokum. 1971 areosphaeridium eaton, p. 357, 358. synopsis. chorate gonyaulacinean cysts with subspherical central body. acavate. processes mesotabular, solid, commonly fibroid, expanded clypeate and fenestrate to reticulate distally. the number of cingular processes is variable. archaeopyle apical, with formula a(1 –4´), operculum free. remarks. the synopsis is based on the emendation of areosphaeridium by stover & williams (1995, p. 100). following fensome et al. (2007), areosphaeridium is considered an areoligeracean cyst related to enneadocysta. areosphaeridium diktyoplokum (klumpp 1953) eaton 1971 (plate 2, fig. 12) 1953 hystrichosphaeridium diktyoplokum klumpp, p. 392, plate 18, figs 3–7 (not plate 18, figs 8–10, which are now cordosphaeridium latum). 1963a cordosphaeridium diktyoplokum (klumpp) – eisenack, p. 262. 1971 areosphaeridium diktyoplokum (klumpp) – eaton, p. 358, 359. age. lo: latest priabonian. genus atopodinium drugg 1978 type. drugg 1978, plate 1, fig. 1, as atopodinium pro statum. 1978 atopodinium drugg, p. 62. 1981 maghrebinia below, p. 22. 1987 bejuia stover & williams, p. 37. remarks. masure (1991, p. 64) demonstrated that ato podinium has a gonyaulacacean tabulation and that the archaeopyle of the type is apical, with the formula a1 –4´ , and with precingular accessory sutures; the operculum is attached. atopodinium cf. haromense thomas & cox 1988 (plate 2, fig. 13) cf. 1988 atopodinium haromense thomas & cox, p. 319–321, 323, plate 1, figs 1–6; text-fig. 4. age. lo: late campanian. remarks. the cretaceous forms found in the present study tend to have slightly more pronounced ornamentation than that of the jurassic holotype of atopodinium haromense (thomas & cox 1988, plate 1, figs 1, 2). genus axiodinium williams, damassa, fensome & guerstein in fensome et al. 2009 type. williams & downie 1966b, plate 18, fig. 1, as wetzeliella articulata; now axiodinium prearticulatum. 2009 axiodinium williams, damassa, fensome & guerstein in fensome et al., p. 16. remarks. both axiodinium and apectodinium are wetzelielloidean genera with an equiepeliform archaeopyle (wil liams et al. 2015), but axiodinium is clearly cavate and apectodinium does not have a clear or consistent separation of the endophragm and periphragm. axiodinium augustum (harland 1979b) williams et al. 2015 (plate 2, figs 14, 15) 1979b wetzeliella (apectodinium) augusta harland, p. 63, plate 2, figs 13–15. 1981 apectodinium augustum (harland) – lentin & williams, p. 14. 2015 axiodinium augustum (harland) – williams et al., p. 301. age. lo: basal ypresian. genus batiacasphaera drugg 1970 type. drugg 1970, fig. 6a–b, as batiacasphaera compta. 1970 batiacasphaera drugg, p. 813. remarks. we adhere to the synopsis presented in fensome et al. (2009, p. 17), which specifies that the apical archaeopyle has an outline lacking or with weakly developed accessory sutures between precingular plates and a free operculum. batiacasphaera micropapillata stover 1977 (plate 2, fig. 16) 1977 batiacasphaera micropapillata stover, p. 73, plate 1, figs 7, 8. bulletin36.qxp_bulletin 36 19/12/16 13.39 side 29 30 age. lo (not confirmed in present study) gelasian on scotian margin (fensome et al. 2009). not plotted. remarks. this is a species of batiacasphaera with a granulate to microreticulate ornamentation. genus batioladinium brideaux 1975 type. alberti 1961, plate 5, fig. 2, as broomea jaegeri. 1975 batioladinium brideaux, p. 124. 1975 necrobroomea wiggins, p. 111. synopsis. an elongate pareodiniacean cyst, drawn out into one apical and two generally equal antapical horns. acavate. wall consisting of autophragm. archaeopyle apical, with formula a(1–2´), formed from the loss of the two apical plates. remarks. the synopsis is based on the interpretation of the archaeopyle presented in wharton (1988, text-fig. 4.13) and fensome et al. (1993, p. 77, 78). the antapical horns, though equal on a single specimen, are of variable length from specimen to specimen. batioladinium jaegeri (alberti 1961) brideaux 1975 (plate 2, fig. 17) 1961 broomea jaergeri alberti, p. 26, plate 5, figs 1–7. 1975 batioladinium jaegeri (alberti) – brideaux, p. 1240. 1975 necrobroomea jaegeri (alberti) – wiggins, p. 111. 1980 imbatodinium jaegeri (alberti) – dörhöfer & davies, p. 37. 1981 pseudoceratium hansgochtii lentin & williams, p. 236. age. lo: campanian. remarks. most records of this species are from the early cretaceous, but it appears to have an extended range up to the campanian in western and northern north america (e.g. harker et al. 1990). genus callaiosphaeridium davey & williams 1966a type. deflandre & courteville 1939, plate 4, fig. 1, as hystrichosphaeridium asymmetricum. 1966a callaiosphaeridium davey & williams, p. 103. 1967 hexasphaera clarke & verdier, p. 42; name illegitimate. remarks. callaiosphaeridium is a chorate gonyaulacacean genus that has large, tubular mesotabular or gonal processes on the cingulum and may have slender gonal processes elsewhere. the archaeopyle is epicystal, with the formula (a1–4´ p1–6´´) and a simple operculum. callaiosphaeridium asymmetricum (deflandre & courteville 1939) davey & williams 1966a (plate 2, fig. 18) 1939 hystrichosphaeridium asymmetricum deflandre & courteville, p. 100, 101, plate 4, figs 1, 2. 1966a callaiosphaeridium asymmetricum (deflandre & courteville) – davey & williams, p. 104. 1967 hexasphaera asymmetrica (deflandre & courteville) – clarke & verdier, p. 43; combination illegitimate. age. lo: campanian. genus cannosphaeropsis wetzel 1933a type. wetzel 1933a, plate 3, fig. 9a, b, as canno sphae r opsis utinensis. 1932 cannosphaeropsis wetzel, p. 136; name not validly published. 1933a cannosphaeropsis wetzel, p. 6. synopsis. chorate gonyaulacacean (gonyaulacoidean) cysts with a subspherical to ellipsoidal central body. tabulation on the central body delineated by gonal processes, which can be restricted to the cingular plates, and sometimes also intergonal processes, which distally are united by trabecula that also reflect the tabulation. the gonal processes are trifurcate distally, the intergonal processes are bifurcate distally. archaeopyle precingular, with formula p3´´; operculum free. cannosphaeropsis passio de verteuil & norris 1996 (plate 3, fig. 5) 1996 cannosphaeropsis passio de verteuil & norris, p. 130, 132, 134, 136, plate 7, figs 1–8; plate 8, figs 1–6; plate 17, figs 1, 3–5; text-fig. 33. age. lo: serravallian (based on piasecki 2003). bulletin36.qxp_bulletin 36 19/12/16 13.39 side 30 31 remarks. de verteuil & norris (1996) described canno sphaeropsis passio as having cingular processes only, even though the trabecula delineate the complete tabulation. genus cerebrocysta bujak in bujak et al. 1980 type. bujak et al. 1980, plate 13, figs 4–5, as cere bro cysta bartonensis. 1980 cerebrocysta bujak in bujak et al., p. 42. remarks. the synopsis for cerebrocysta provided by fen some et al. (2009, p. 18) covers all the salient details of the morphology of this genus. these authors also noted that it is difficult to separate cerebrocysta from pyxidinopsis. the only distinctions are in the possible variability in the number of precingular plates lost in archaeopyle formation in cerebrocysta and perhaps different wall structures. regardless, we follow fensome et al. (2009) in assigning cenozoic taxa with the general morphology of cere brocysta and pyxidinopsis to cerebrocysta. cerebrocysta bartonensis bujak in bujak et al. 1980 (plate 2, figs 19, 20) 1980 cerebrocysta bartonensis bujak in bujak et al., p. 42, plate 13, figs 4–7. age. lo: bartonian. cerebrocysta magna bujak 1994 (plate 3, fig. 6) 1994 cerebrocysta magna bujak, p. 121, plate 2, figs 10, 11. age. lo: lutetian. genus cerodinium vozzhennikova 1963 type. vozzhennikova 1963, text-fig. 9, as cerodinium sibiricum. 1963 cerodinium vozzhennikova, p. 181. 1963 ceratiopsis vozzhennikova, p. 181; name illegitimate. remarks. fensome et al. (2009, p. 18) noted that cero dinium is characterised by having an isodeltaform to isothetaform 2a plate, a free perioperculum and a symmetrical antapex. the endoarchaeopyle may involve all three anterior intercalary plates. at present it does not seem practicable to separate into distinct species forms that lose the 2a plate only from those that lose the 1a, 2a and 3a plates individually from the endocyst. however, such a separation seems to have some stratigraphic value as forms occurring in younger rocks all seem to lose the 2a plate only; earlier forms may lose all three anterior intercalaries. cerodinium diebelii (alberti 1959) lentin & williams 1987 (plate 3, figs 1–4) 1959 deflandrea diebelii alberti, p. 99, 100, plate 9, figs 18–21. 1967 ceratiopsis diebelii (alberti) – vozzhennikova, p. 159; combination illegitimate. 1987 cerodinium diebelii (alberti) – lentin & williams, p. 114. age. lo: latest danian. cerodinium glabrum (gocht 1969) fensome et al. 2009 (plate 3, figs 7, 8) 1969 deflandrea speciosa forma glabra gocht, p. 10, text-fig. 3. 1973 deflandrea speciosa subsp. glabra (gocht) – lentin & williams, p. 45. 1977a ceratiopsis speciosa subsp. glabra (gocht) – lentin & williams, p. 21 (combination illegitimate). 1987 cerodinium speciosum subsp. glabrum (gocht) – lentin & williams, p. 115. 2009 cerodinium glabrum (gocht) – fensome et al, p. 19. age. lo: latest thanetian (late paleocene). cerodinium kangiliense nøhr-hansen & heilmann-clausen 2001 (plate 3, figs 9, 10) 2001 nøhr-hansen & heilmann-clausen, p. 158, 160, 162–164, fig. 4, nos 1–9, fig. 5, nos 1–9. age. lo: selandian. cerodinium speciosum (alberti 1959) lentin & williams 1987 (plate 3, fig. 13) bulletin36.qxp_bulletin 36 19/12/16 13.39 side 31 32 1959 deflandrea speciosa alberti, p. 97, plate 9, figs 12, 13. 1977a ceratiopsis speciosa (alberti) – lentin & williams, p. 21; combination illegitimate. 1987 cerodinium speciosum (alberti) – lentin &williams, p. 115. age. lo: latest thanetian. cerodinium striatum (drugg 1967) lentin & williams 1987 (plate 3, figs 11, 12) 1967 deflandrea striata drugg, p. 18, plate 2, figs 13, 14. 1977a ceratiopsis striata (drugg) – lentin & williams, p. 21; combination illegitimate. 1987 cerodinium striatum (drugg) – lentin & williams, p. 115. age. lo thanetian. not plotted. remarks. cerodinium striatum has a distinct development of more or less parallel folds. the folds should not be confused with the plications of cerodinium diebelii (which is distinctly elongate) or the linearly aligned granules or denticles in cerodinium speciosum. genus charlesdowniea lentin & vozzhennikova 1989 emend. williams et al. 2015 type. williams & downie 1966b, plate 18, fig. 8, textfig. 47, as wetzeliella coleothrypta, and bujak et al. 1980, plate 12, figs 7, 8, as kisselevia coleothrypta. 1989 charlesdowniea lentin & vozzhennikova, p. 225, 227. remarks. the emended diagnosis of williams et al. (2015) is followed; this emphasises the archaeopyle type and the nature of the pericystal ornamentation. the archaeopyle of charlesdowniea is equiepeliform and the processes are intratabular and sometimes also penitabular. distally, the processes on individual plates are united by membranes. charlesdowniea coleothrypta (williams & downie 1966b) lentin & vozzhennikova 1989 (plate 4, fig. 9) 1966b wetzeliella coleothrypta williams & downie, p. 185, 186, plate 18, figs 8, 9; text-fig. 47. 1976 kisselevia coleothrypta (williams & downie) – lentin & williams, p. 136. 1989 charlesdowniea coleothrypta (williams & downie) – lentin & vozzhennikova, p. 225. age. lo: ypresian. not plotted. remarks. in the labrador margin samples, charles dow n iea coleothrypta has its lo just below that of scale no dinium scalenum, but is not common. genus chatangiella vozzhennikova 1967 emend. nov. type. vozzhennikova 1967, plate 56, fig. 1; plate 57, fig. 1, as chatangiella niiga. 1967 chatangiella vozzhennikova, p. 128, 129. 1967 australiella vozzhennikova, p. 129, 130. 1967 cooksoniella vozzhennikova, p. 183, 184. emended description. peridiniacean (deflandreoid) cysts that are proximate, dorsoventrally compressed and peridinioid in outline, with epicystal ‘shoulders’ and asymmetrical antapex, the left side being larger. cir cumcavate to, generally, bicavate. surface generally atabulate to weakly paratabulate, but cingulum clearly reflected by ridges, commonly serrated, reflecting the positions of preand postcingular plates. periarchaeopyle intercalary, with formula i2a, operculum free or attached; plate 2a is latito iso-omegaform hexa; endoarchaeopyle with for mula i2a or i1a +2a +3a. remarks. fensome et al. (2009, p. 19) provided a synopsis for chatangiella that encompasses all its salient morphologic features. as part of their synopsis, these authors stated: “periarchaeopyle intercalary, with formula i2a, operculum free or attached; plate 2a is isodelta form to (typically) iso-omegaform hexa; endoarchaeopyle with formula i2a or i1a + 2a + 3a.” we disagree with the synopsis of fensome et al. (2009) in two important respects. firstly, only taxa with omegaform 2a plates should be included in chatangiella; and, secondly, taxa with 2a plates that are lati-omegaform must be included since the type of chatangiella (vozzhennikova 1967, plate 56, fig. 1; plate 57, fig. 1) has such an archaeopyle. thus it is proposed here that a characteristic feature of chatangiella be the possession of a latior iso-omegaform hexa 2a plate. in their emendation of chatangiella, lentin & wil liams (1976, p. 51, 52) emphasised the omegaform shape of the pericystal 2a plate and its common attachment bulletin36.qxp_bulletin 36 19/12/16 13.39 side 32 33 posteriorly. unfortunately, the critical shape and relationship of this plate have been overlooked in many subsequent papers. to correct this concern, the omega form shape of the operculum is re-emphasised in the emended description above. although epicystal shoulders, a distinctive characteristic of this genus, appear to almost always be related to the presence of an omegaform periarchaeopyle, which feature came first is open to debate. as noted under alterbidinium, the morphological differences used to separate the deflandreoid genera are in large part unsatisfactory. this criticism applies to chatangiella: to bring stability to the genus, we propose that any peridiniacean forms with a partite cingulum that do not have an omegaform periarchaeopyle should be re-assigned. chatangiella decorosa (mcintyre 1975) lentin & williams 1976 1975 deflandrea decorosa mcintyre, p. 63, 64, plate 2, figs 1–4. 1976 chatangiella decorosa (mcintyre) – lentin & williams, p. 54. age. lo: early campanian. remarks. mcintyre (1975, p. 64) noted that chatangiella decorosa differs from chatangiella ditissima in being larger. the size of chatangiella ditissima overlaps that of chatangiella tripartita. chatangiella decorosa also has more pustules than chatangiella ditissima. chatangiella madura lentin & williams 1976 (plate 3, figs 14, 15) 1970b deflandrea manumii cookson & eisenack, p. 141, 142, plate 11, figs 10, 11. 1976 chatangiella manumii (cookson & eisenack) – lentin & williams, p. 54; name illegitimate. 1976 chatangiella madura (cookson & eisenack) – lentin & williams, p. 54. age. lo: campanian. chatangiella tripartita (cookson & eisenack 1960a) lentin & williams 1976 (plate 3, figs 16, 20) 1960a deflandrea tripartita cookson & eisenack, p. 2, 3, plate 1, fig. 10. 1967 australiella tripartita (cookson & eisenack) – vozzhennikova, p. 134, 135. 1976 chatangiella tripartita (cookson & eisenack) – lentin & williams, p. 55. age. lo: campanian (late cretaceous). not plotted. remarks. there is considerable confusion concerning the morphology of chatangiella tripartita and the related form chatangiella victoriensis. both are surprisingly large. according to cookson & eisenack (1960a, p. 3), the pericyst of chatangiella tripartita ranges from 100 to 120 μm in length and 59 to 71 μm in width. cookson & manum (1964) stated that in chatangiella victoriensis, the pericyst was 76 to 116 μm in length and 49–73 μm in width. comparison of the sizes of the above two species with one of the higher latitude (arctic) species of chatangiella, chatangiella ditissima, shows that there is some overlap, especially in width. mcintyre (1975, p. 63) noted that the pericyst in chatangiella ditissima could be 115–150 μm long and 60–90 μm wide. the similar but even larger, higher latitude species chatangiella decorosa has a pericyst that can be 130–175 μm long and 80–110 μm wide. our conclusion from comparison of several chatan giella species is that size is not a reliable characteristic when separating species. but size has some significance: for example specimens of chatangiella found on the grand banks of newfoundland are generally much smaller than those from the high arctic. gigantism is a common feature of life in higher latitudes, so it is no surprise to find it among dinoflagellates. differences in the expression of the cingulum and the surface ornamentation of the pericyst are perhaps more diagnostic. cookson & manum (1964, p. 521, 522) noted that the cingulum of the holotype of chatangiella tripartita is denoted by two pairs of ill-defined short, offset, low parallel ridges on the ventral surface and a fold-like line on the dorsal surface. the ridges are not serrate or denticulate, a feature of many specimens placed in chatangiella tripartita. regarding the cingulum of chatangiella victoriensis, cookson & manum (1964) stated that it “is bordered by conspicuous ridges or by linearly arranged wart-like thickenings of varying size and shape”. mcintyre (1975, p. 62, 63) described chatangiella ditissima as having a cingulum delineated by “discontinuous ridges consisting of rows of pustules that may join to form narrow grooves”. that author recognised seven and five ridges on the anterior and posterior margins respectively of the cingulum, which represent the difference in number of plates between the preand bulletin36.qxp_bulletin 36 19/12/16 13.39 side 33 34 postcingular series. these two series of plates can also have penitabular pustules mirroring the plate outlines. cookson & manum (1964, p. 522) gave a thickness of 1.0–1.7 μm for the periphragm (as the thecal wall) of chatangiella victoriensis and noted that it is “ornamented with fairly evenly scattered rod-like projections c. 0.5–1.5 μm long; in surface view the rods appear as dots usually between 0.5 and 1.0 μm in diameter but a few smaller and larger ones are usually present”. in cha tangiella tripartita, the periphragm is finely granular. mcintyre (1975) described the periphragm of cha tangiella ditissima as smooth, except for the pustules delineating the tabulation and occasional pits. what is the conclusion to be drawn from the above and from our observations on the specimens from the offshore wells? the most compelling is the extreme variability in the size of the taxa and the degree of expressions of tabulation. in this paper, all the ‘smaller’ forms are assigned to chatangiella tripartita if they lack pustules other than on the cingulum. we would like to include those smaller forms with pustules defining the preand postcingular plates in a new species but do not have enough specimens to do that here. genus chiropteridium gocht 1960 emend. nov. type. gocht 1960, plate 17, fig. 1, as chiropteridium lobospinosum. 1959 galea maier, p. 305; name illegitimate. 1960 chiropteridium gocht, p. 221. emended diagnosis. areoligeracean cysts that are proximochorate to chorate, with subrounded to lenticular central body that may be asymmetrical antapically. cavate or acavate; if cavate, with cavation mainly restricted to scalloped marginal wings; when acavate there may be marginal wings and/or processes. midventral and mid-dorsal areas devoid of processes or having processes of a reduced size. surface atabulate or with intratabular processes. archaeopyle apical, with formula a(1–4´) and free operculum; sulcal notch offset to the left. remarks. chiropteridium has not been emended previously. however, stover & evitt (1978) provided a synopsis and modified description and fensome et al. (2009, p. 21) gave a synopsis. in their synopsis of chiropteridium, stover & evitt (1978, p. 27) stated: “cysts skolochorate, body lenticular; processes isolated or connected proximally and absent or greatly reduced in size and number on ventral surface; archaeopyle apical, type ta; parasulcal notch offset.” in their modified description, the same authors noted that processes could be isolated or partly connected in longitudinal rows, with the connections being proximal. also they stated that processes are absent or reduced in size and numbers ventrally. fensome et al. (2009, p. 21) were more specific in providing the following synopsis for chiropteridium: “areoligeracean cysts that are proximochorate to chorate, with lenticular central body that may be asymmetrical antapically. cavate or acavate; if cavate, with cavation mainly restricted marginally to ‘wings’ and scalloped; if acavate, marginate wings formed from longitudinal crests or from longitudinally taeniate processes. surface atabulate. archaeopyle apical, with formula a(1–4´), operculum free; sulcal notch offset to the left.” fensome et al. (2009, p. 21) stated that “chirop teridium is characterised by lateral extensions or ‘wings’ that may be cavate; these wings are typically scalloped to varying degrees or formed by taeniate processes. in membranophoridium, processes are absent and the ‘wings’ are continuous pericoelar sacs.” although correct for species included to date in chiropteridium, forms are recorded here that possess only processes that are not connected proximally or along their length and which denote the tabulation. these are accommodated in the above emended diagnosis for chiropteridium. chiropteridium galea (maier 1959) sarjeant 1983 (plate 3, figs 17–19) 1959 galea galea maier, p. 306, plate 29, fig. 4; text-fig. 2. 1959 galea levis maier, p. 308, plate 30, figs 1, 2. 1959 galea mespilana maier, p. 306, 307, plate 29, figs 5, 6. 1960 chiropteridium dispersum gocht, p. 227, plate 18, figs 1–16; text-figs 16–27. 1961 membranophoridium multispinatum gerlach, p. 203, 204, plate 29, fig. 5. 1961 membranophoridium partispinatum gerlach, p. 201, plate 29, fig. 6. 1963 chiropteridium partispinatum (gerlach) – brosius, p. 48. 1964 baltisphaeridium leve (maier) – sarjeant, p. 176. 1964 baltisphaeridium mespilanum (maier) – sarjeant, p. 176. 1969 cleistosphaeridium leve (maier) – davey et al., p. 16. 1973 chiropteridium mespilanum (maier) – lentin & williams, p. 26. bulletin36.qxp_bulletin 36 19/12/16 13.39 side 34 35 1975 hystrichosphaeridium mespilanum (maier) – eisenack & kjellström, p. 233, 234. 1983 chiropteridium galea (maier) – sarjeant, p. 108. age. lo: chattian. chiropteridium gilbertii sp. nov. (plate 4, figs 1–4) holotype. plate 4, fig. 4, from a cuttings sample at 1600–1610 m in gilbert f-53, gsc type collection no.137976, sample p39466, slide 01, co-ordinates 17.4 × 106.0, england finder r36/3. overall length 83 μm, width 87 μm; central body, length 57 μm, width 52 μm; maximum length of processes 20 μm, width varies from less than 1 to 10 μm. the age determined for the sample from which the holotype was recovered is basal bartonian. etymology. the epithet is derived from the name of the gilbert f-53 offshore exploration well, in which the species is abundant. diagnosis. a species of chiropteridium in which the processes or process complexes delineate the tabulation and there are no interconnections between processes representing adjacent plates. description. the processes delineate the tabulation, though none are interconnected. processes on the midventral and mid-dorsal surfaces are reduced in width or absent: this is especially true of the 3´´ process that, if present, is very slender. commonly, processes are basally trumpet-shaped, then tubular or parallel-sided. pro cesses are also of variable size depending on which plate series they are reflecting, with cingular processes being invariably slender, precingular processes broader, and postcingular processes being broader still, and the antapical process being broadest of all. many of the processes are perforate. size. overall maximum diameter 87 μm; maximum diameter of central body 57 μm; maximum length of processes 25 μm, maximum width 17 μm; seven specimens measured. age. lo: bartonian. remarks. in chiropteridium gilbertii, the central body, which is granulate, commonly has a prominent antapical protuberance. all of the processes, which are in tratabular, have a circular cross-section proximally and distally are closed and commonly branch into short bifurcations or aculeae. proximally, the processes tend to be conical then become tubular about halfway along their length before flaring distally. the processes are often fibrous and/or perforate, especially the antapical process, which sometimes subdivides to form a process complex. recognition of the antapical process, if the archaeopyle is not obvious, is facilitated by the extreme width and the perforations, which may form arches in the process wall. cingular processes are invariably slender, some less than 1 μm wide, rarely exceed four in number, and are restricted to the ambital region of the central body. apical processes show variation in width on different specimens. precingular processes are always narrower than the postcingular processes. although the sulcal notch is offset, this is not always obvious. chiropteridium gilbertii differs from chirop te ri dium galea in having processes or process complexes that are restricted to individual plates. licracysta semicirculata has processes that are restricted to the ambitus and tend to form arcuate complexes. the processes of chiro p teridium gilbertii vary in size according to the plate they represent, and are often fibrous and/or perforate. chirop teridium conispinum also has membranous pro cesses, but these are restricted to the dorsal surface, with two linear membranes running apically–antapically on the ventral surface. genus chlamydophorella cookson & eisenack 1958 type. cookson & eisenack 1958, plate 11, fig. 1, as chlamydophorella nyei. 1958 chlamydophorella cookson & eisenack, p. 56. remarks. as fensome et al. (2009, p. 21) noted in their synopsis for chlamydophorella, the genus is characterised by being proximate, holocavate, with short, solid processes forming buttresses between the autophragm and ectophragm, an apical horn formed from the ectophragm, and an apical archaeopyle. the tabulation is gonyaulacacean, but whether it shows neutral or dextral torsion is not known. sepispinula differs from chlamydophorella in lacking an ectophragm. chlamydophorella nyei cookson & eisenack 1958 (plate 4, fig. 5) bulletin36.qxp_bulletin 36 19/12/16 13.39 side 35 36 1958 chlamydophorella nyei cookson & eisenack, p. 56, plate 11, figs 1–3. 1970b chlamydophorella apiculata cookson & eisenack, p. 150, 151, plate 13, fig. 3. 1970b chlamydophorella lagena cookson & eisenack, p. 151, plate 13, fig. 4. age. lo: coniacian (late cretaceous). chlamydophorella cf. nyei cookson & eisenack 1958 (plate 4, figs 6–8) age. lo : campanian. not plotted. remarks. this form includes specimens similar to chla mydophorella nyei but which lack an evident apical horn. genus chytroeisphaeridia (sarjeant 1962) downie & sarjeant 1965 type. sarjeant 1962, plate 70, fig. 13, as leiosphaeridia subgenus chytroeisphaeridia chytroeides. 1962 leiosphaeridia subgenus chytroeisphaeridia sarjeant, p. 492. 1965 chytroeisphaeridia downie & sarjeant, p. 102. chytroeisphaeridia hadra sp. nov. (plate 4, figs 17, 18) holotype. plate 4, fig. 18 from a cuttings sample at 3120–3140 m in roberval k-92, gsc type collection no. 137902, sample p17728, slide 01, co-ordinates 19.5 × 109.8, england finder u40/1. overall length 79 μm, width 81 μm; wall thickness 5 μm, archaeopyle length 37 μm, width 33 μm. the age determined for the sample from which the holotype was recovered is late cretaceous (the specimen is caved). etymology. the epithet is from the greek hadros meaning well-developed, bulky, stout, strong, great, in reference to the thick, sturdy wall. diagnosis. a species of chytroeisphaeridia with a thick wall that has a subdued rugulate, sometimes lightly striated surface. the archaeopyle is large, apparently resulting from the loss of the 3´´ plate only. size. overall length 69–82 μm; width 64–81 μm. six specimens measured. age. lo: bartonian. remarks. some specimens of chytroeisphaeridia hadra have parallel striations on the wall but these are irregular and discontinuous. the archaeopyle is large, suggesting that more than a single plate may be involved. genus cleistosphaeridium davey et al. 1966 type. davey et al. 1966, plate 10, fig. 7, as cleisto sphaeridium diversispinosum. 1966 cleistosphaeridium davey et al., p. 166. remarks. we concur with the retention and emendation of cleistosphaeridium as proposed by eaton et al. (2001, p. 176, 177) and the synopsis provided by fensome et al. (2009, p. 22). following fensome et al. (2007, p. 408), cleistosphaeridium is considered to be areoligeracean because of the asymmetry in some spe cimens of the antapex and sulcal notch, and also because it forms a morphological plexus with other areoligeracean genera, including glaphyrocysta, enne adocysta, licracysta and cooksonidium. we include in the genus species in which adjacent processes are medially and distally interconnected. cleistosphaeridium diversispinosum davey et al. 1966 (plate 4, figs 19, 20) 1966 cleistosphaeridium diversispinosum davey et al., p. 167, plate 10, fig. 7. 1993 systematophora diversispinosa (davey et al.) – islam, p. 88. age. lo: serravallian. remarks. following fensome et al. (2009, p. 22), cleisto sphaeridium diversispinosum and cleistosphaeridium ancyr eum are not considered synonyms. cleistosphaeridium elegantulum sp. nov. (plate 4, figs 10–12) holotype. plate 4, fig. 12, from a cuttings sample at 2286.03 to 2295.17 m in karlsefni a-13, gsc type collection no. 138033, sample p39600, slide 01, co-ordinates 17.9 × 101.7, england finder s44.2. central body maximum diameter 50 μm, maximum length of processes about 25 μm. the age determined for the sample from which the holotype was recovered is lutetian–bartonian. bulletin36.qxp_bulletin 36 19/12/16 13.39 side 36 37 etymology. from the latin elegantulus (very fine), in reference to the narrow, long processes. diagnosis. a species of cleistosphaeridium with nume rous long, fine, flexible, unconnected processes, some to many of which are dolabrate distally. the length of many of the processes equals about one half of the central body diameter. proximal ridges are absent to weakly developed. size. central body diameter 30‒57 μm; maximum length of processes up to half maximum diameter of central body; three specimens measured. age. lo: lutetian‒bartonian. not plotted. remarks. this species differs from other species of cleistosphaeridium in its long, fine, flexible processes. cleistosphaeridium palmatum sp. nov. (plate 4, figs 13–16) holotype. plate 4, figs 13, 14 from a cuttings sample at 2450–2460 m in roberval k-92: gsc type collection no. 137897, sample p17706, slide 01, co-ordinates 6.9 × 100.6, england finder g30/0‒2. maximum overall diameter 87 μm, maximum diameter of central body 50 μm, maximum length of processes 25 μm. the age determined for the sample from which the holotype was recovered is early ypresian. etymology. the epithet is from the latin palmatus, meaning marked or shaped like the palm of the hand, in reference to the shape of the distal terminations of the processes. diagnosis. a species of cleistosphaeridium in which the solid processes are of irregular width, varying from 1–7 μm and distally slender to splayed, with some adjacent processes interconnected distally, commonly so that they form arches. description. the processes are of irregular width, distally slender to splayed, with some adjacent processes interconnected distally. they are predominantly dolobrate, but some are bifid. the wall of the central body is generally less than 1 μm thick. archaeopyle apical, with formula a(1–4´), operculum attached or free. other aspects of the tabulation cannot be determined, as specimens possess more than one process per plate, sometimes in complexes. where present, the complexes appear to occur on the precingular, postcingular and antapical plates. size. maximum diameter of central body 48 μm, length of processes 10–25 μm, four specimens measured. age. lo: late ypresian. remarks. the processes of specimens belonging to this species show similarities to those found in adnato sphaeridium, enneadocysta, licracysta and other species of cleistosphaeridium. although the distal branches of pro cesses can sometimes be interconnected, the general absence of trabecula uniting processes precludes assignment of the species to adnatosphaeridium. cleistospha er idium palmatum differs from species of enneadocysta because the number of processes per plate (where dis cernible as being related to tabulation) invariably exceeds one, and distally processes are dolobrate rather than licrate or clypeate. moreover, the antapex of cleis to s phae r idium palmatum is not characterised by two pro cesses (see fensome et al. 2007, p. 394). cleistospha eridium diversispinosum differs in not having branched processes in which the branches often meet along their length or distally and in not having spatulate endings. specimens of the species cleistosphaeridium polypetellum are slender and are not spatulate or branched distally; species of licracysta have proximal and ventral surfaces on which processes are absent or reduced. cleistosphaeridium polypetellum (islam 1983b) stover & williams 1995 (plate 5, figs 1, 2) 1983b areosphaeridium polypetellum islam, p. 82, 84, plate 2, figs 1–6. 1995 cleistosphaeridium polypetellum (islam) – stover & williams, p. 102. age. lo: ypresian. remarks. some connections occur between adjacent processes, differentiating this species from cleistos phae r idium diversispinosum and cleistosphaeridium an cyreum. processes in cleistosphaeridium polypetellum are less variable and the species lacks the thick processes of clei sto s phaeridium palmatum. distally, most processes of cleisto sphaer idium polypetellum are strongly dola brate and sometimes licrate. however, a few simple bifid processes occur. bulletin36.qxp_bulletin 36 19/12/16 13.39 side 37 38 genus cordosphaeridium eisenack 1963a type. klumpp 1953, plate 18, figs 1, 2, as hystricho sphaeridium inodes. 1963a cordosphaeridium eisenack, p. 261. 1981 tityrosphaeridium sarjeant, p. 120. remarks. we concur with the synopsis for cordos phae r idium provided by fensome et al. (2009, p. 22). the characteristic features of the genus are the spheroidal central body, the precingular (p3´´) archaeopyle and the fibrous wall and processes. the processes, which are commonly cylindrical and restricted to one per plate, are generally of approximately equal length but can show considerable variation in width. cordosphaeridium cantharellus (brosius 1963) gocht 1969 (plate 5, fig. 7) 1963 hystrichosphaeridium cantharellus brosius, p. 40, 41, plate 6, fig. 1; text-fig. 2, nos 11a–c. 1969 cordosphaeridium cantharellus (brosius) – gocht, p. 45. 1981 tityrosphaeridium cantharellus (brosius) – sarjeant, p. 120. age. lo: burdigalian. remarks. as in the holotype, the process width in the labrador margin specimens of cordosphaeridium can tharellus shows considerable variation. the species was retained in cordosphaeridium by edwards (2001, p. g19). cordosphaeridium delimurum fensome et al. 2009 (plate 5, figs 3, 4) 2009 cordosphaeridium delimurum fensome et al., p. 23, plate 2, figs l–q. age. lo: early lutetian. not plotted. remarks. as its name implies, cordosphaeridium delimurum differs from cordosphaeridium inodes and cor dosphaeridium gracile in having a much thinner central body wall. unlike cordosphaeridium fibrospinosum, cor do sphaeridium delimurum has solid rather than perfo rate process walls. cordosphaeridium fibrospinosum davey & williams 1966a (plate 5, figs 5, 6) 1966a cordosphaeridium fibrospinosum davey & williams, p. 86, plate 5, fig. 5. 1966a cordosphaeridium exilimurum davey & williams, p. 87, 88, plate 11, fig. 2. 1970 achomosphaera valianta sah et al., p. 145, plate 1, figs 8, 9. 1978 cordosphaeridium valiantum (sah et al.) – stover & evitt, p. 147. 1981 hystrichosphaerina? exilimura (davey &williams) – sarjeant, p. 122. 1981 emmetrocysta? fibrospinosa (davey & williams) – sarjeant, p. 123. 1986 tityrosphaeridium? exilimurum (davey & williams) – jain & garg, p. 120. 1986 tityrosphaeridium? fibrospinosum (davey & williams) – jain & garg, p. 121. age. lo: late rupelian. not plotted. remarks. as noted above, cordosphaeridium fibrospino sum has processes with fibrous, perforate walls. we agree with fensome et al. (2009) that cordosphaeridium exilimurum and its long-recognised synonym achomo sphaera valianta cannot be meaningfully differentiated from cordosphaeridium fibrospinosum. although this spe cies has an oligocene lo, at least in offshore eastern canada, it tends to be most common in the paleocene. cordosphaeridium funiculatum morgenroth 1966a (plate 5, fig. 8) 1966a cordosphaeridium funiculatum morgenroth, p. 22, 23, plate 6, figs 2–3. 1981 tityrosphaeridium funiculatum (morgenroth) – sarjeant, p. 121. age. lo: priabonian. cordosphaeridium gracile (eisenack 1954) davey & williams 1966a (plate 5, fig. 9) 1954 hystrichosphaeridium inodes subsp. gracile eisenack, p. 66, plate 8, fig. 17; plate 10, figs 3–8; plate 12, figs 7, 21. 1963a cordosphaeridium inodes subsp. gracile (eisenack) – eisenack, p. 261. bulletin36.qxp_bulletin 36 19/12/16 13.39 side 38 39 1966a cordosphaeridium gracile (eisenack) – davey & williams, p. 84. 1981 tityrosphaeridium gracile (eisenack) – sarjeant, p. 121. age. lo: early lutetian. cordosphaeridium inodes (klumpp 1953) eisenack 1963a (plate 5, fig. 10) 1953 hystrichosphaeridium inodes klumpp, p. 391, plate 18, figs 1, 2. 1963a cordosphaeridium inodes (klumpp) – eisenack, p. 261. age. lo: consistent occurrence early lutetian. not plotted. remarks. the holotype of cordosphaeridium inodes (klumpp 1953, plate 18, figs 1, 2) shows considerable variation in the width of the processes, some being remarkably similar to those of cordosphaeridium gracile. it is this variability in width, however, that is the distinguishing characteristic of cordosphaeridium inodes. genus cribroperidinium neale & sarjeant 1962 type. neale & sarjeant 1962, plate 19, fig. 4, text-fig. 3a, b, as cribroperidinium sepimentum. 1962 cribroperidinium neale & sarjeant, p. 443. 1968 acanthaulax sarjeant, p. 227. 1978 millioudodinium stover & evitt, p. 173. 1984b meristaulax sarjeant, p. 160. 1988 meristaulax brenner, p. 65. synopsis. gonyaulacacean (cribroperidinioid) cysts that are proximate, with spheroidal to more commonly ovoidal central body, usually surmounted by an apical horn. acavate or cornucavate. tabulation strongly delineated by sutural and penitabular ornamentation. archae o pyle pre cingular, with formula p3´´, operculum free. cribroperidinium giuseppei (morgenroth 1966a) helenes 1984 1966a gonyaulax giuseppei morgenroth: 5, plate 2, figs 3–6. 1969 gonyaulacysta giuseppei (morgenroth) – sarjeant, p. 9. 1978 millioudodinium? giuseppei (morgenroth) – stover & evitt, p. 174. 1982 rhynchodiniopsis? giuseppei (morgenroth) – sarjeant, p. 36. 1984 cribroperidinium giuseppei (morgenroth) – helenes, p. 121. age. local acme: priabonian. remarks. cribroperidinium giuseppei is a species of cribro peridinium with a small, button-like apical horn, a spon gy wall, distinctly reflected paratabulation and a low, coarse reticulum superimposed on the paraplates. this includes many forms assigned in the literature to cri broperidinium tenuitabulatum, but the latter species is now considered to be a junior synonym of apteodinium australiense. helenes (1984, p. 122) noted that the sutural ridges in cribroperidinium giuseppei are smooth and that the penitabular ridges are more commonly found on the hypocyst. genus cyclonephelium deflandre & cookson 1955 type. deflandre & cookson 1955, plate 2, fig. 12, as cyclonephelium compactum. 1955 cyclonephelium deflandre & cookson, p. 285. 1961 circulodinium alberti, p. 28. remarks. we follow the synopsis for this genus pro posed by fensome et al. (2009, p. 24), and agree with those authors in considering circulodinium to be a junior synonym of cyclonephelium. cyclonephelium distinctum (deflandre & cookson 1955) jansonius 1986 (plate 6, fig. 10) 1955 cyclonephelium distinctum deflandre & cookson, p. 285, 286, plate 2, fig. 14; textfigs 47, 48. 1961 circulodinium hirtellum alberti, p. 28, 29, plate 4, fig. 20. 1969 canningia hirtella (alberti) – millioud, p. 425. 1978 cyclonephelium hirtellum (alberti) – davey, p. 894. 1986 circulodinium distinctum (deflandre & cookson) – jansonius, p. 204. age. lo: latest maastrichtian. bulletin36.qxp_bulletin 36 19/12/16 13.39 side 39 40 genus dapsilidinium bujak et al. 1980 type. davey & williams 1966a, plate 4, fig. 10, as poly sphaeridium pastielsii. 1980 dapsilidinium bujak et al., p. 27, 28. remarks. for this genus, the concept of fensome et al. (2009, p. 26) is followed. dapsilidinium pseudocolligerum (stover 1977) bujak et al. 1980 (plate 5, fig. 13) 1977 polysphaeridium pseudocolligerum stover, p. 74, 75, plate 1, figs 14–19. 1980 dapsilidinium pseudocolligerum (stover) – bujak et al., p. 28. age. lo: tortonian. remarks. as fensome et al. (2009) observed, dapsili dinium pastielsii and dapsilidinium pseudocolligerum differ primarily in the proximal morphology of the processes. in dapsilidinium pastielsii, the processes are initially broad before tapering gradually, whereas in dapsilidinium pseu docolligerum they are more or less cylindrical proximally, tapering gradually to the distal opening. dapsilidinium pseudoinsertum sp. nov. (plate 5, figs 11, 12) holotype. plate 5, fig. 12 from a cuttings sample at 2825‒2835 m in rut h-11, gsc type collection no. 137957, sample p39388, slide 01, co-ordinates 19.6 × 105.1, england finder u48/1‒2. central body maximum diameter 32 μm; length of processes up to about 15 μm. the age determined for the sample from which the holotype was recovered is ypresian. etymology. the epithet is in reference to the similarity of this species to hystrichokolpoma? incertum michoux 1985. diagnosis. a species of dapsilidinium in which some, but not all, of the processes are significantly broader than others, though the position of the broader processes appears not to be consistent. size. central body maximum diameter 30‒32 μm; length of processes up to about 15 μm; two specimens measured. age. lo: late lutetian. remarks. this species is distinctive in having a more or less bimodal variation in process widths, with a few being distinctly broader than the others. the position and number of broad processes vary from specimen to specimen. one or more of the broader processes on each spe cimen may be bifurcate distally down to about midlength. in hystrichokolpoma? incertum, the larger processes are consistently the precingulars and the antapical. in other species of dapsilidinium, such as dap silidinium pseudocolligerum and dapsilidinium simplex, the processes are more or less uniformly developed. in species of diphyes, only the antapical process is relatively large. dapsilidinium simplex (white 1842) bujak et al. 1980 (plate 5, fig. 14) 1842 xanthidium tubiferum var. simplex white, p. 38, plate 4, fig. 10. 1946 hystrichosphaeridium simplex (white) – deflandre, card 934. 1969 polysphaeridium? simplex (white) – davey & williams, p. 7. 1980 dapsilidinium simplex (white) – bujak et al., p. 28. age. lo: bartonian. not plotted. genus deflandrea eisenack 1938 type. eisenack 1938, text-fig. 6, as deflandrea phosphoritica. 1938 deflandrea eisenack, p. 187. remarks. deflandrea is a peridiniacean cyst that is characterised by its latideltaform intercalary (i2a) archaeopyle. the broad shape is a consistently striking feature of the periarchaeo pyle, but not always of the endoarchaeopyle. deflandrea borealis sp. nov. (plate 5, figs 17–20) holotype. plate 5, fig. 19 from a cuttings sample at 3095‒3105 m in rut h-11, gsc type collection no. 137964, sample p39397, slide 01, co-ordinates 18.9 × 98.6, england finder t41/2. length (including horns) 57 μm, width 55 μm. the age determined for the sample from which the holotype was recovered is tha netian (late paleocene). etymology. the epithet is from the latin borealis (northern) in reference to the northern occurrence of at least the type material of this species. bulletin36.qxp_bulletin 36 19/12/16 13.39 side 40 41 diagnosis. a relatively small, squat and generally rounded species of deflandrea with a scabrate to granulate wall and a latiform archaeopyle, the operculum of which com monly remains attached posteriorly. size. length (including horns) 46‒59 μm, width 46‒55 μm; three specimens measured. age. lo: latest priabonian. deflandrea denticulata alberti 1959 (plate 5, figs 15, 16) 1959 deflandrea denticulata alberti, p. 102, 103, text-fig. 1. age. lo: ypresian. not plotted. remarks. deflandrea denticulata (alberti 1959, text-fig. 1) is unusual for a species of this genus in having a pericyst with long slender apical and antapical horns. the periphragm is covered with small, slender spines. al though the archaeopyle appears to be latideltaform, it is impossible to be definite without examining the holotype. alberti (1959) recorded deflandrea denticulata from lower eocene sediments of volgograd, russia, the oebisfelde borehole in germany, and possibly from belgium. his observations fit the predicted lo in this study for a taxon with a morphology intermediate between cerodinium and deflandrea. deflandrea galeata (lejeune-carpentier 1942) lentin & williams 1973 (plate 6, fig. 1) 1942 peridinium galeata lejeune-carpentier, p. b186–b188, figs 15–20. 1973 deflandrea galeata (lejeune-carpentier) – lentin & williams, p. 41. age. lo: maastrichtian. remarks. the line-drawing of the dorsal view of the holotype of deflandrea galeata (lejeune-carpentier 1942, fig. 15) clearly shows a latideltaform hexa 2a archae o pyle. this interpretation is confirmed in le jeune-carpentier & sarjeant (1981, p. 18, 19, who refer to the archaeopyle as “single-plate intercalary (type i/i) of broad-hexa type and formed by the loss of paraplate 2a”. in every other respect, the species has the typical morphology of cerodinium. it is unusual to find forms with a latideltaform 2a in rocks of this age. deflandrea majae (schiøler 1993) comb. nov. (plate 6, fig. 2) 1993 isabelidinium majae schiøler, p. 108, 110, plate 1, figs 1–6; text-fig. 4a, b. age. lo: latest maastrichtian. remarks. schiøler (1993) noted that this species has a latideltaform archaeopyle with a transverse archaeopyle index (tai) considerably higher than 0.5. schiøler (1993, p. 110) stated that although these characters are “… typical of the genus deflandrea … as the new species lacks any signs of a paracingulum, referral to the latter genus is precluded.” in our view, archaeopyle shape is critical in diagnosing deflandrea and similar genera and that the presence or absence of a cingulum is not significant. therefore, this species is transferred herein to deflandrea. however, m. pearce (personal communication 2015) has pointed out that the archaeo pyles of specimens otherwise attributable to this species show a wide variation in archaeopyle shape. for example he has observed specimens of deflan drea majae with stenodeltaform archaeopyles; we would recommend that such forms be included in another genus. as noted under alterbidinium, a de tailed re-evaluation of archaeo pyle shapes in deflan drea and similar genera in relation to the taxonomy of the group is clearly needed but beyond the scope of the present work. deflandrea oebisfeldensis alberti 1959 (plate 6, fig. 3) 1959 deflandrea oebisfeldensis alberti, p. 95, 96, plate 8, figs 10–13. age. lo: early ypresian. remarks. the antapical margin of the pericyst in de flandrea oebisfeldensis is broad, with the left and right antapical horns being at the intersection with the posterior lateral sides and thus broadly separated. deflandrea phosphoritica eisenack 1938 (plate 6, fig. 4) 1938 deflandrea phosphoritica eisenack, p. 187, text-fig. 6. bulletin36.qxp_bulletin 36 19/12/16 13.39 side 41 42 1965b deflandrea granulosa cookson & eisenack, p. 122, plate 11, figs 8, 9. 1965 deflandrea heterophlycta forma pusulosa rozen, p. 293, 294, plate 1, figs 3, 4; text-fig. 2. 1966 deflandrea menendezii pöthe de baldis, p. 223, plate 2, fig. a. 1973 deflandrea heterophlycta subsp. pusulosa (rozen) – lentin & williams, p. 41. age. lo: latest chattian. genus dinogymnium evitt et al. 1967 type. evitt et al. 1967, plate 1, figs 21–23, plate 2, fig. 5, text-figs 16–18, as dinogymnium acuminatum. 1967 dinogymnium evitt et al., p. 4–8. remarks. lentin & vozzhennikova (1990) subdivided the dinogymnioids into genera that are separated on overall outline, relative size of the episome and hyposome, and surface ornamentation. as a consequence, these authors transferred several species previously in cluded in dinogymnium into the new genera. fensome et al. (2009, p. 27) provided a synopsis for dino gym nium that is followed here. dinogymnium longicorne (vozzhennikova 1967) harland 1973 (plate 6, fig. 11) 1967 gymnodinium longicorne vozzhennikova, p. 46, plate 1, fig. 8; plate 3, fig. 6; plate 4, figs 6a, b, 7. 1967 gymnodinium curvatum vozzhennikova, p. 43, plate 1, figs 10–12; plate 4, figs 2, 3. 1973 dinogymnium longicorne (vozzhennikova) – harland, p. 678. age. lo: early campanian. remarks. dinogymnium longicorne is elongate, with an episome that is considerably longer than the hyposome. according to lentin & vozzhennikova (1990, p. 19), the length varies between 62 and 91 μm, the width between 21 and 38 μm. genus diphyes cookson 1965 nom. cons. type. deflandre & cookson 1955, plate 7, fig. 3, as hystrichosphaeridium colligerum. 1965 diphyes cookson, p. 85; name illegitimate. 1970 lingulasphaera drugg, p. 817. 2000 diphyes cookson nom. prop. cons. harris & fensome, p. 281, 282. remarks. the conservation proposal of the name diphyes cookson (harris & fensome 2000) was ratified at the 2005 botanical congress. we agree with the synopsis of diphyes in fensome et al. (2009, p. 28), with one exception regarding the processes. some specimens show an intratabular organisation, with about four processes on each preand postcingular plate. diphyes brevispinum bujak 1994 (plate 6, figs 5, 6) 1994 diphyes brevispinum bujak, p. 121, 123, plate 2, figs 4–6. age. lo: ypresian. remarks. diphyes brevispinum has short conical to subconical, rarely tapering processes and an inflated antapical process that is similar to that of diphyes ficusoides. diphyes colligerum (deflandre & cookson 1955) cookson 1965 (plate 6, fig. 12) 1955 hystrichosphaeridium colligerum deflandre & cookson, p. 278, 279, plate 7, fig. 3. 1965 baltisphaeridium colligerum (deflandre & cookson) – downie & sarjeant, p. 88. 1994 diphyes pseudoficusoides bujak, p. 123, 125, plate 2, figs 2, 3. age. lo: lutetian. remarks. following fensome et al. (2009, p. 30), di phyes pseudocolligerum should be considered a junior synonym of diphyes colligerum because the size of the antapical process is similar in both holotypes. diphyes ficusoides islam 1983a (plate 6, figs 7, 8) 1983a diphyes ficusoides islam, p. 338, plate 2, figs 8, 9. age. lo: middle lutetian. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 42 43 genus disphaerogena wetzel 1933a type. wetzel 1933a, plate 4, fig. 34, as disphaerogena carposphaeropsis. 1933a disphaerogena wetzel, p. 51. 1976 cyclapophysis benson, p. 192. 1981 plethysyrinx sarjeant, p. 106. remarks. the synopsis of disphaerogena provided in fensome et al. (2009, p. 30) is accepted, including that cy clapophysis is a taxonomic junior synonym of the genus. disphaerogena carposphaeropsis wetzel 1933a (plate 6, fig. 9) 1933a disphaerogena carposphaeropsis wetzel, p. 51, plate 4, fig. 34. 1976 cyclapophysis monmouthensis benson, p. 183, plate 1, figs 9–12; plate 2, fig. 1. age. lo: latest maastrichtian. remarks. when sarjeant (1985a, p. 141, 142) emended the diagnosis of disphaerogena carposphaeropsis, he considered cyclapophysis monmouthensis to be a taxonomic junior synonym of the species. genus eatonicysta stover & evitt 1978 type. morgenroth 1966a, plate 3, fig. 11, as canno sphae ropsis ursulae. 1978 eatonicysta stover & evitt, p. 41. synopsis. gonyaulacacean (leptodinioid) cysts that are chorate, with a spheroidal central body. holocavate. central body bearing 17 to 23 hollow or solid, fibroid mesotabular processes, which are distally connected by a fenestrate to reticulate to irregular open-mesh, net-like ectophragm. archaeopyle apical, with formula a(1–4´), operculum free. description. the holocavate cyst has an autophragm bearing mesotabular processes with expanded distal extremities that merge to form a perforate membranous to reticulate to trabeculate ectophragm. the size of the mesh or perforations shows considerable variation. a leptodinioid tabulation of 3–4´, 6´´´´, 0–6c, 5´´´, 1´´´´, 0s is indicated by the processes. based on process size, plate 1´´ would be wider than 6´´. remarks. eatonicysta is characterised by a membranous ectophragm that may be reticulate or broken down to form a trabeculate network. eatonicysta furensis (heilmann-clausen in heil mannclausen & costa 1989) stover & williams 1995 (plate 6, fig. 13) 1989 eatonicysta ursulae subsp. furensis heilmannclausen in heilmann-clausen & costa, p. 466, plate 11, figs 3, 5, 7. 1995 eatonicysta furensis (heilmann-clausen in heilmann-clausen & costa) – stover & williams, p. 104. age. lo: late ypresian. remarks. eatonicysta furensis has much shorter, usually broader, funnel-shaped processes than eatonicysta ursulae. also, the processes grade imperceptibly into the ecto phragm, which is divided into areas reflecting the indivi dual plates, rather than forming a continuous network. eatonicysta ursulae (morgenroth 1966a) stover & evitt 1978 (plate 6, fig. 14) 1966a eatonicysta ursulae morgenroth, p. 20, plate 3, figs 11, 12. 1966a membranilarnacia reticulata williams & downie, p. 220, 221, plate 24, figs 4, 6; text-fig. 59. 1967 membranilarnacia dictyophora agelopoulos, p. 49, 50, plate 12, figs 3, 4, 6. 1969 membranilarnacia ursulae (morgenroth) – de coninck, p. 43. 1978 eatonicysta ursulae (morgenroth) – stover & evitt, p. 41. age. lo: earliest lutetian. remarks. williams & downie (1966a) recorded two variants of eatonicysta ursulae (as membranilarnacia re ticulata, a synonym of eatonicysta ursulae) from the ypresian london clay from southern england. one lacked cingular processes, the other had four cingular processes. whether or not this variation is stratigraphically significant remains to be confirmed. genus enneadocysta stover & williams 1995 type. gerlach 1961, plate 28, fig. 14, as baltisphae r idium pectiniforme. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 43 44 1994 enneadocysta stover & williams in bujak, p. 119; name not validly published. 1995 enneadocysta stover & williams, p. 108, 109. 2007 enneadocysta stover & williams – emend. fensome et al., p. 394. remarks. we follow the emendation of fensome et al. (2007, p. 394) in our concept for this genus and agree that it is areoligeracean. enneadocysta magna fensome et al. 2007 (plate 7, figs 1, 2) 2007 enneadocysta magna fensome et al., p. 394, 396, plate 1, figs 1–20; plate 2, figs 1–19; text-figs 5a, b, 6a–e. age. lo: latest rupelian. remarks. labrador margin specimens of enneadocysta magna can have process clusters on individual plates. genus eocladopyxis morgenroth 1966a type. morgenroth 1966a, plate 3, figs 2, 3, as eoclado pyxis peniculata. 1966a eocladopyxis morgenroth, p. 7. remarks. the synopsis provided by fensome et al. (2009, p. 31) takes into account that eocladopyxis is a goniodomacean (pyrodinioid) cyst and identifies all the plates involved in the formation of the archaeopyle. eocladopyxis peniculata morgenroth 1966a (plate 7, fig. 3) 1966a eocladopyxis peniculata morgenroth, p. 7, 8, plate 3, figs 2, 3. age. lo: late ypresian. genus evittosphaerula manum 1979 emend. damassa 1997 type. manum 1979, plate 2, figs 3, 4, as evittosphae rula paratabulata. 1979 evittosphaerula manum, p. 242, 243. 1997 evittosphaerula manum – emend. damassa, p. 161–163. evittosphaerula? foraminosa sp. nov. (plate 6, figs 15–20) holotype. plate 6, figs 19, 20 from a cuttings sample at 2130–2140 m in north leif i-05, gsc type collection no. 138159, sample yd17600, slide 03, co-ordinates 44.9 × 15.5, england finder m43/0. pericyst length 82.5 μm, width 90 μm. the age determined for the sample from which the holotype was recovered is ypresian. etymology. from the latin foraminosus, meaning ‘full of holes’. description. a species of gonyaulacalean cysts in which only broad strips of membrane representing the sutures are preserved. the tabulation appears to be gonio doma ceans, with a five-sided antapical plate reflecting a quinqueform hypocystal tabulation. at the apex, sutural strips come together to form a short apical horn. size. diameter 68–90 μm; four specimens in dorsoventral orientation measured, but excluding the holotype as it is oriented apically–antapically. age. lo: ypresian. remarks. this species has a very distinctive structure, represented by strips of membrane reflecting the sutures only; the internal area of each plate is represented by a hole. thus the species is reminiscent of the late oligo cene to early miocene species evittospaherula paratabulata, but differs superficially in having an apical horn, an apparently much narrower cingulum, and broader sutural membranes. more fundamentally, one specimen (plate 6, fig. 18) appears to have a five-sided ant apical plate, suggesting a goniodomacean affinity (fen some et al. 1993). the tabulation described by manum (1979) for evittosphaerula paratabulata, the type of the genus, is clearly gonyaulacacean. as this new species is strikingly similar, albeit perhaps superficially, to manum’s species, and there is not sufficient material to describe its tabulation in full, it is questionably assigned to evittosphaerula. the new species is also strikingly similar to both hapsocysta susanae, described by duxbury (2002) from the albian of the central north sea, and chaenos phaerula magnifica, described by damassa (1997) from the late oligocene of the norwegian sea. however, the tabulation of both those species is clearly sexiform, and neither has horns. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 44 45 genus fibrocysta stover & evitt 1978 type. cookson & eisenack 1965b, plate 16, fig. 8, as cordosphaeridium bipolare. 1978 fibrocysta stover & evitt, p. 155. synopsis. gonyaulacacean (cribroperidinioid) cysts that are chorate, with a longitudinally elongate ovoidal central body, with a protrusion at the apical and antapical poles. acavate or cornucavate, with a fibrous wall. processes numerous, nontabular to indistinctly tabular; they are hollow and fibrous or solid and generally of uniform size. archaeopyle precingular, with formula p3´´, operculum free. remarks. in wall structure and archaeopyle type, fibro cysta is very similar to turbiosphaera. however, processes in tur bio sphaera are wider, especially apically and antapically, and in turbiosphaera, the cingulum is marked by a membrane. the remarkable similarities be tween the two genera suggest that fibrocysta, like turbiosphaera, has a cribroperidinioid tabulation. fibrocysta bipolaris (cookson & eisenack 1965a) stover & evitt 1978 (plate 7, fig. 4) 1965a cordosphaeridium bipolare cookson & eisenack, p. 135, plate 16, figs 7, 8. 1969 lanternosphaeridium bipolare (cookson & eisenack) – de coninck, p. 38. 1969c amphorosphaeridium bipolare (cookson & eisenack) – davey, p. 35. 1978 fibrocysta bipolaris (cookson & eisenack) – stover & evitt, p. 155. age. local acme early ypresian. remarks. the cingulum of the holotype of fibrocysta bipo laris (cookson & eisenack 1965a, plate 16, fig. 8) appears to be delineated by a single row of processes. genus gillinia cookson & eisenack 1960a type. cookson & eisenack 1960a, plate 3, fig. 4, as gillinia hymenophora. 1960a gillinia cookson & eisenack, p. 11, 12. synopsis. small, proximate, slightly elongate cysts with two membranous wings, one on each side of the apical archaeopyle. tabulation partially delineated by ridges, which commonly define the cingulum and a sulcal area that is considerably broader posteriorly. remarks. following fensome et al. (1993, p. 73), gil linia probably has a cladopyxiacean tabulation. gillinia hymenophora cookson & eisenack 1960a (plate 7, fig. 12) 1960a gillinia hymenophora cookson & eisenack, p. 12, plate 3, figs 4–6; text-fig. 5. age. lo: late campanian. genus ginginodinium cookson & eisenack 1960a type. cookson & eisenack 1960a, plate 2, fig. 9, as gin gi no dinium spinulosum. 1960a ginginodinium cookson & eisenack, p. 7. remarks. in their emendation of ginginodinium, len tin & williams (1976, p. 95, 96) noted that the archaeopyle is compound, involving the three anterior intercalary plates and three of the precingular plates (3´´–5´´). however, the archaeopyle may be formed from loss of only the three anterior intercalary plates. there may also be a series of successive stages, until all three intercalaries are lost and three of the precingulars remain attached to the main cyst solely along the cingular margin. this variability in archaeopyle type often makes assignment of species to the genus difficult. ginginodinium? flexidentatum sp. nov. (plate 7, figs 5–11) holotype. plate 7, fig. 11, from a cuttings sample at 1815–1825 m in bjarni o-82, gsc type collection no. 138070, sample p39715, slide 01, co-ordinates 3.6 × 102.8, england finder c32/4. pericyst length 84 μm, width 73 μm, endocyst length 61 μm, width 67 μm, length of processes up to 2.5 μm. the age determined for the sample from which the holotype was recovered is early ypresian. etymology. the epithet is from the latin flexibilis, meaning bendable and dentatus, meaning toothed or pointed. diagnosis. a species of ginginodinium with one or more lateral horns or bulges, a flexible folded wall, and no clearly demarked cingulum. the ornament is variable but typically denticulate. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 45 46 description. pericyst outline pentagonal, with a welldeveloped apical and two antapical horns, one of which is slightly shorter than the other, and two lateral protuberances. generally cornucavate but can, in part or whole, be circumcavate. the endocyst is pentagonal. pericyst ornamentation varies from finely perforate to verrucate to bearing processes. the archaeopyle in volves the loss of the 2a plate, which remains attached to the 4´´ plate, and sometimes the partial detachment of the 1a and 3a plates. one margin of the cingulum is commonly indicated by a fold. size. pericyst length 61–87 μm (mean 73 μm), pericyst width 56–73 μm (mean 63 μm), endocyst length 43– 61 μm (mean 51 μm), width 49– 67 μm (mean 53 μm), process length up to about 2.5 μm. age. lo: late ypresian. remarks. the shapes of the pericyst and endocyst are relatively stable in ginginodinium? flexidentatum, but there are differences between specimens in degree of cavation, which can range from cornucavate to narrowly circumcavate. there are also variations in the shape of the horns, especially the apical: this is generally rounded distally but may be acuminate. the antapical horns are acuminate to rounded distally; the lateral horns are primarily just protuberances. pericyst ornamentation is variable, sometimes even on the same specimen. the periphragm can be perforated and ornamented with small verrucae and processes, the latter of variable length. processes are slender and distally bifid. this species is included in ginginodinium only provisionally because of the uncertainty over the archaeopyle. genus glaphyrocysta stover & evitt 1978 type. cookson 1965, plate 11, fig. 4, as cyclonephe lium retiintextum. 1978 glaphyrocysta stover & evitt, p. 49, 50. remarks. the generic concept of glaphyrocysta, as ex pressed in the emendation by fensome et al. (2009, p. 32), is followed here. the genus is sometimes abundant in middle eocene sections in labrador margin wells. glaphyrocysta divaricata (williams & downie 1966a) stover & evitt 1978 (plate 7, figs 15, 16) 1966a cyclonephelium divaricatum williams & dow nie, p. 223, 224, plate 25, fig. 1; text-fig. 60. 1978 glaphyrocysta divaricata (williams & downie) – stover & evitt, p. 50. age. lo: earliest lutetian. see also remarks. remarks. a species of glaphyrocysta with numerous pro cesses variously interconnected along their length, but with a preponderance of distal free ends. in glaphyro cysta divaricata, unlike in glaphyrocysta ordinata and gla phyrocysta retiintexta, the processes do not form distinct complexes. although this taxon has a lutetian lo, it is most abundant in the paleocene and is often very common to dominant in assemblages from that epoch. glaphyrocysta exuberans (deflandre & cookson 1955 ex eaton 1976) stover & evitt 1978 (plate 7, fig. 14) 1955 cyclonephelium exuberans deflandre & cookson, p. 255; name not validly published. 1976 cyclonephelium exuberans deflandre & cookson ex eaton, p. 255, 256. 1978 glaphyrocysta exuberans (deflandre & cookson ex eaton) – stover & evitt, p. 50. age. lo: priabonian. remarks. glaphyrocysta exuberans has a well-developed, often perforate, marginate ectophragm, supported by solid, slender processes. glaphyrocysta retiintexta (cookson 1965) stover & evitt 1978 (plate 7, fig. 17) 1965 cyclonephelium retiintextum cookson, p. 88, plate 11, fig. 4. 1978 glaphyrocysta retiintexta (cookson) – stover & evitt, p. 50. age. lo: priabonian. remarks. glaphyrocysta retiintexta shows an ambital development of distal trabecula but has minimal membrane development. the general absence of membranes and the distal connections between plate complexes distinguish this species from glaphyrocysta intricata. fensome et al. (2009, p. 34) noted that the morphology of glaphyrocysta retiintexta was very similar to that shown in the drawing bulletin36.qxp_bulletin 36 19/12/16 13.40 side 46 47 of the holotype of glaphyrocysta pastielsii (deflandre & cookson 1955) stover & evitt 1978 by pastiels (1948, plate 5, fig. 15). but fensome et al. (2009) also stated that photographs of the holotype of glaphyrocysta pastielsii ap peared to possess a morphology more like that of gla phy rocysta divaricata. accordingly they recommended restricting the name glaphyrocysta pastielsii to the holotype, a proposal followed here. glaphyrocysta texta (bujak 1976) stover & evitt 1978 (plate 8, fig. 1) 1976 cyclonephelium texta bujak, p. 110, plate 3, figs 6–11; text-fig. 3g, h. 1978 glaphyrocysta texta (bujak) – stover & evitt, p. 50. age. lo: priabonian. remarks. glaphyrocysta texta has distinctive process com plexes on most of the preand postcingular plates, and apparently one broader complex on the single antapical plate. the contabular complexes consist of slender in tratabular processes, distally united to form clypeate platforms that have ragged margins. the 3´´ and 3´´´ plates sometimes have a single process rather than a process complex. adjacent process complexes are connected distally by trabecula. glaphyrocysta vicina (eaton 1976) stover & evitt 1978 (plate 8, fig. 2) 1976 cyclonephelium vicinum eaton, p. 260, 261, plate 8, figs 4, 5; text-fig. 13. 1978 glaphyrocysta vicina (eaton) – stover & evitt, p. 50. age. lo: lutetian. remarks. glaphyrocysta vicina has a marginal, perforate pericoel, with the periphragm remaining close to the endophragm. both features are unusual in species of gla phy rocysta. species of membranophoridium lack pro cesses supporting the periphragm. genus habibacysta head et al. 1989 type. head et al. 1989, plate 4, figs 1, 2, 5, 6, as habi bacysta tectata. 1989 habibacysta head et al., p. 457, 458. synopsis. proximate spheroidal gonyaulacean cysts. ata bulate. autophragm bearing short columns that distally are sometimes united by an entire, perforate or reticulate layer. archaeopyle precingular, with formula p3´´, operculum free. remarks. head (1994, text-fig. 3) showed how the nature of the autophragm differentiates habibacysta from bitectatodinium, filisphaera and tectatodinium. all four genera are proximate, atabulate gonyaulacacean cysts with precingular archaeopyles. habibacysta tectata head et al. 1989 (plate 7, fig. 13) 1989 habibacysta tectata head et al., p. 458, plate 4, figs 1–6, 9, 10. age. lo: earliest zanclean. genus hapsocysta davey 1979 type. eisenack & cookson 1960, plate 3, fig. 6, as can nosphaeropsis peridictya. 1979 hapsocysta davey, p. 556. synopsis. a chorate to camocavate gonyaulacacean (go ny au la coidean) cyst, with a subrounded to ovoidal central body, which is surrounded by a periphragm that may be trabeculate, forming an open network, or filled by a per forate membrane. tabulation indicated by the trabe cula, which on membranous taxa occur as ridges. archaeo pyle precingular, with formula p3´´, operculum free. remarks. fensome et al. (1993, p. 89) considered hap socysta to have a cribroperidinioid tabulation. however, nøhr-hansen (1993, p. 71, 72) showed that the taxon, which he named hapsocysta? benteae, had gonyaulacoidean tabulation. nøhr-hansen (1993) provisionally included the species in hapsocysta, because of the presence of the thin-walled periphragm. hapsocysta? benteae nøhr-hansen 1993 (plate 8, fig. 12) 1993 hapsocysta? benteae nøhr-hansen, p. 71, 72, plate 25, figs 11, 12; text-figs 10a, b, 11a, b. age. lo: cenomanian. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 47 48 genus heteraulacacysta drugg & loeblich jr. 1967 type. drugg & loeblich jr. 1967, plate 1, fig. 8a–c, as heteraulacacysta campanula. 1967 heteraulacacysta drugg & loeblich jr., p. 183. remarks. the synopsis for heteraulacacysta by fen some et al. (2009, p. 35) is followed. heteraulacacysta porosa bujak in bujak et al. 1980 (plate 8, fig. 16) 1980 heteraulacacysta porosa bujak in bujak et al., p. 62, plate 15, figs 10–13; text-fig. 14b, c. age. lo: priabonian. remarks. considerable uncertainty has prevailed over the distinction between heteraulacacysta leptalea eaton 1976 and heteraulacacysta porosa. in his diagnosis for heteraulacacysta leptalea, eaton (1976, p. 305) stated: “circular fenestrations frequently developed in the proximal area of the cingular crests, along with fine elongate fenestrations aligned at right-angles to the margin of the cyst body.” under ‘remarks’, eaton commented that small perforations occurred in the periphragm, comparable to those in the proximal area of the cingular crests, and commonly gave a punctate appearance to the wall. in the succeeding sentence, eaton (1976, p. 305) differentiated heteraulacacysta leptalea from heteraulacacysta campanula on the basis of its “…frequently punctuate rather than positively ornamented surface to the cyst body, and in exhibiting circular and elongate fenestrations in the cingular crests.” bujak in bujak et al. (1980, p. 62) differentiated heteraulacacysta porosa “… from all other described species of heteraulacacysta by its perforate periphragm.” the illustrations of the holotype and another specimen of heteraulacacysta leptalea in eaton (1976, plate 21, figs 1, 2) clearly show perforations. consequently, the only consistent difference appears to be the presence of elongate perforations on the cingular crests in hete rau lacacysta leptalea. we do not know, however, whether this feature is stratigraphically significant. genus heterosphaeridium cookson & eisenack 1968 type. cookson & eisenack 1968, text-fig. 4h, as he tero sphae ridium conjunctum. 1968 heterosphaeridium cookson & eisenack, p 115. remarks. in their synopsis for heterosphaeridium, fen some et al. (2009, p. 35) recognised that it is an areoligeracean cyst with a spheroidal to broadly ovoidal central body. yun hyesu (1981, p. 45, 46) provided an emend ed diagnosis that broadened the circumscription mainly to include forms with hollow, open processes as well as solid processes. some of the specimens of heteros phae r idium difficile observed in the present study have pro cesses that are perforated or circular (annulate) process complexes that can be connected at various locations along their length. heterosphaeridium bellii radmacher et al. 2014 (plate 8, figs 3, 4) 2014 heterosphaeridium bellii radmacher et al., p. 31–33, plate 1, figs 1–9. age. lo: late campanian. remarks. heterosphaeridium bellii differs from hetero sphaeridium heteracanthum in having dolabrate pro cesses that are not branching. the processes in het e ro s phae r idium bellii may be connected proximally but are never branched along their length. radmacher et al. (2014) considered the lo of heterosphaeridium bellii to be late campanian to early maastrichtian in the southwestern barents sea. heterosphaeridium difficile (manum & cookson 1964) ioannides 1986 (plate 8, fig. 8) 1964 hystrichosphaeridium difficile manum & cookson, p. 12–14, plate 3, figs 1–3, 7. 1986 heterosphaeridium difficile (manum & cookson) – ioannides, p. 24. age. lo: early santonian. remarks. labrador margin specimens of heteros phae r idium difficile show extreme variation in the nature of the processes. some are perforate along their length or arched so that the process walls merge distally. others, as noted under ‘remarks’ for the genus, have process complexes on the preand postcingular plates that are connected by circular proximal membranes, as in systematophora. in the latter specimens, the cingular plates have linear process complexes. in some specimens, the apical plates com bulletin36.qxp_bulletin 36 19/12/16 13.40 side 48 49 monly show a characteristic x-shaped crest as observed by m. pearce (personal communication 2015) and also appear to have annulate process complexes. the antapical plate is marked by an annulate complex that is broader than those on the preand postcingular plates. the wall of the central body is microreticulate. genus histiocysta davey 1969a type. davey 1969a, plate 1, fig. 5, text-fig. 14a, b, as histiocysta palla. 1969a histiocysta davey, p. 138. remarks. this genus is similar to corrudinium in its shape and ornamentation, but tends to be smaller and has an apical archaeopyle. histiocysta palla davey 1969a (plate 7, figs 18–20) 1969a histiocysta palla davey, p. 138–140, plate 1, figs 5, 6; text-fig. 14a, b. age. lo: late campanian. genus homotryblium davey & williams 1966a type. davey & williams 1966a, plate 12, fig. 5, as ho mo tryblium tenuispinosum. 1966a homotryblium davey & williams, p. 100. remarks. the synopsis provided by fensome et al. (2009, p. 35, 36) covers all the key morphological attributes of homotryblium. homotryblium abbreviatum eaton 1976 (plate 8, figs 9, 10) 1976 homotryblium abbreviatum eaton, p. 267, 268, plate 10, figs 2–4. age. lo: late ypresian. not plotted. remarks. as in the type material, the labrador margin specimens of homotryblium abbreviatum have short and wide processes, but the surface ornamentation of the central body can be smooth as well as granulate. this species is not as common as homotryblium tenuis pi no sum in the samples studied here. homotryblium tenuispinosum davey & williams 1966a (plate 8, figs 5–7) 1966a homotryblium tenuispinosum davey & williams, p. 101, 102, plate 4, fig. 11; plate 12, figs 1, 5, 7; text-fig. 21. 1966a homotryblium pallidum davey & williams, p. 102, 103, plate 12, figs 4, 6; text-fig. 22. age. lo: bartonian; peak: latest ypresian. genus hystrichokolpoma klumpp 1953 type. klumpp 1953, plate 17, figs 3, 5a, as hystri cho kolpoma cinctum. 1953 hystrichokolpoma klumpp, p. 388. remarks. fensome et al. (2009, p. 36) provided a de tailed synopsis for hystrichokolpoma, which stipulated among other morphologic attributes, that the genus must have cingular processes. hystrichokolpoma cinctum klumpp 1953 (plate 8, fig. 11) 1953 hystrichokolpoma cinctum klumpp, p. 389, plate 17, figs 3, 4, 5a–d. age. lo: burdigalean according to williams et al. (2004), but we consider that the age range of this species needs to be better constrained. not plotted. remarks. a species of hystrichokolpoma in which the preand postcingular processes fill plates and have small tubular extensions. there are generally two processes per cingular paraplate. hystrichokolpoma globulus michoux 1985 (plate 8, figs 13–15) 1985 hystrichokolpoma globulus michoux, p. 143, plate 1, figs 1–4, 12; text-fig. 2a, b. age. lo: late ypresian. not plotted. remarks. michoux (1985, p. 143) compared hystricho kol poma globulus to hystrichokolpoma cinctum klumpp 1953, from which it differs by having a much shorter antapical process relative to the length of the cyst and preand postcingular processes that do not branch into bulletin36.qxp_bulletin 36 19/12/16 13.40 side 49 50 tubules distally. also, the cingular and sulcal processes are conical and there are three to five on each cingular plate. the labrador margin specimens of hystricho kolpoma globulus differ from the type material in not having conical cingular plates and in there being only three or less on each plate. genus hystrichosphaeridium deflandre 1937 type. ehrenberg 1838, plate 1, fig. 16, as xanthidium tubiferum. 1937 hystrichosphaeridium deflandre, p. 68. remarks. we concur with the synopsis provided by fensome et al. (2009, p. 38) except to note that individual processes proximally do not cover most of the underlying plate and that the cingular and sulcal processes are slender. the apical processes also tend to be slender, especially the first and fourth. variation in process dimensions in hystrichosphaeridium does not approach the strong variation that is distinctive of hystrichokolpoma. hystrichosphaeridium quadratum sp. nov. (plate 8, figs 17–19) holotype. plate 8, figs 17, 18, from a cuttings sample at 2770–2780 m in gilbert f-53, gsc type collection no. 137988, sample p39505, slide 01, co-ordinates 11.3 × 99.2, england finder l29/3. diameter of central body 44 μm, length of processes up to about 37 μm. the age determined for the sample from which the holotype was recovered is early ypresian. etymology. from the latin quadratus (four-cornered), in reference to the rectangular distal endings of the processes in this species. diagnosis. a species of hystrichosphaeridium in which the distal process endings are slightly flared and poly gonal, generally rectangular and commonly perforate. the process bases are mesotabular and circular. size. width of central body 36‒48 μm, length of cen tral body (when operculum in place) 46‒63 μm; pro cesses up to 39 μm long. six specimens measured. age. lo: selandian. not plotted. remarks. this species differs from hystrichos phaeri dium tubiferum in having rectangular endings to its processes. it also resembles hystrichokolpoma proprium, but the holotype of that species has rounded polygonal process bases that largely fill the plate (see fauconnier & masure 2004, plate 40, figs 1–3), as is typical of the genus hystrichokolpoma. hystrichos phae ridium salpingophorum differs in lacking perforations in the distal endings of the processes. hystrichosphaeridium tubiferum (ehrenberg 1838) deflandre 1937 (plate 8, fig. 20) 1838 xanthidium tubiferum ehrenberg, plate 1, fig. 16. 1937 hystrichosphaeridium tubiferum (ehrenberg) – deflandre, p. 68. age. lo: lutetian. genus hystrichosphaeropsis deflandre 1935 type. deflandre 1935, plate 8, fig. 11, as hystri cho sphaeropsis ovum. 1935 hystrichosphaeropsis deflandre, p. 232. 1937 hystrichosphaera subgenus hystrichosphaeropsis (deflandre) deflandre, p. 67. remarks. eisenack (1963b, p. 118) retained hystricho sphaeropsis at generic rank. fensome et al. (2009, p. 38) provided a synopsis that allows for the inclusion of circumcavate forms in hystrichosphaeropsis. hystrichosphaeropsis perforata schiøler 1993 (plate 9, fig. 4) 1993 hystrichosphaeropsis perforata schiøler, p. 106, plate 2, figs 4–8; text-fig. 3. age. lo: late maastrichtian. hystrichosphaeropsis quasicribrata (wetzel 1961) gocht 1976 (plate 9, fig. 12) 1961 triblastula quasicribrata wetzel, p. 340, plate 2, fig. 3. 1976 hystrichosphaeropsis quasicribrata (wetzel) – gocht, p. 322. age. lo: late maastrichtian. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 50 51 genus hystrichostrogylon agelopoulos 1964 type. agelopoulos 1967, text-figs 1, 2, as hystri cho stro gylon membraniphorum. 1964 hystrichostrogylon agelopoulos, p. 673, 674. hystrichostrogylon digitus sp. nov. (plate 9, figs 1–3) holotype. plate 9, fig. 3, from a cuttings sample at 2435‒2445 m in rut h-11, gsc type collection no. 137952, sample p39375, slide 01, co-ordinates 15.8 × 105.0, england finder q48/1. length of central body (excluding cavation) 45 μm, width of central body (excluding cavation) 38 μm; overall length of cyst 73 μm, overall width of cyst 72 μm. the age determined for the sample from which the holotype was recovered is lutetian priabonian. etymology. from the latin digitus (finger), in reference to the long extensions at the process endings. the epithet is a noun in apposition. diagnosis. a species of hystrichostrogylon in which the biand trifurcations of the processes constitute long, fine, cylindrical distal branches generally one-third to one half the length of the process stem. the distal branches tend to be perpendicular to the process stem. size. length of central body (excluding cavation) 40‒45 μm, width of central body (excluding cavation) 33‒40 μm; overall length of cyst 73‒82 μm, overall width of cyst 66‒79 μm; 3 specimens measured. age. lo: bartonian‒priabonian. not plotted. remarks. this species differs from other species of hy s tric ho strogylon in having remarkably long, fine branches at the ends of the processes. genus impagidinium stover & evitt 1978 type. cookson & eisenack 1965b, plate 12, figs 5–6, as leptodinium dispertitum. 1978 impagidinium stover & evitt, p. 165, 166. synopsis. proximate gonyaulacaceans with the s-type ven tral organisation; in dorso-ventral view they are subspheroidal to ellipsoidal. tabulation clearly defined by sutural septa or thickenings, although sometimes these features are missing between plates bordering the sulcus or the cingulum. sometimes the septum or thickening between the 3´ and 4´ plates is reduced or absent. cingulum and sulcus clearly defined. archae opyle precingular, with formula p3´´, operculum free. remarks. the inclusion of impagidinium in the sub family gonyaulacoideae by fensome et al. (1993, p. 92) reflects an appreciation of the commonly triangular shape of the 6´´ plate. if a separate reflected plate, the 4´ is generally not in contact with the anterior sulcal plate. impagidinium victorianum (cookson & eisenack 1965b) stover & evitt 1978 (plate 9, figs 9–11, 13, 14) 1965b leptodinium victorianum cookson & eisenack, p. 123, plate 12, figs 8, 9. 1978 impagidinium victorianum (cookson & eisenack) – stover & evitt, p. 166. age. lo: intra late maastrichtian. remarks. cookson & eisenack (1965b, p. 123) noted that impagidinium victorianum described from the late eocene differs from impagidinium dispertitum in its “larger size, spherical form, and the constant absence of the transverse dividing between the upper two plates of the ventral field….” where the anterior ventral side of the specimens in this study can be observed, impag idinium victorianum also lacks the suture between what are purportedly the upper two plates of the sulcus. however, labrador sea specimens of impagidinium victorianum are generally smaller than the type material, which ranges from 80 to 120 μm in length and 80 to 123 μm in width. in contrast, the labrador sea specimens vary from 60 to 71 μm in length and 56 to 80 μm in width. distal and proximal ends of the cingulum are offset by approximately a cingulum width. the sutural features may be ridges or septa but are consistent within an individual specimen. maximum height of the sutural ridges or septa is 5 μm. in some specimens the septa are perforated, the perforations aligned in single rows. whether those forms with perforations should be in cluded in a new species is dependent upon finding more specimens. an intriguing aspect of the occurrence of im pagidinium victorianum in the labrador sea is that its lo is consistently within the maastrichtian. this may reflect changing environments from deep-water to shallower conditions, but this does not explain the absence bulletin36.qxp_bulletin 36 19/12/16 13.40 side 51 52 of the species in deeper water parts of the palaeogene. one explanation advanced by m. pearce (personal com munication 2015) is that impagidinium victorianum is a warmer-water species that migrated south during cooler climatic conditions. if so, it is odd that it did not migrate back during paleocene warming, unless opening of the labrador–baffin seaway had already generated a proto-labrador current. genus impletosphaeridium morgenroth 1966a type. morgenroth 1966a, plate 10, fig. 5, as impleto sphaeridium transfodum. 1966a impletosphaeridium morgenroth, p. 32. 1971 ciliosphaeridium grigorovich, p. 94. 1984c laticavodinium wilson & sarjeant in sarjeant, p. 127. remarks. in their remarks on impletosphaeridium, fen some et al. (2009, p. 38) agreed with islam (1993, p. 84, 85) that the archaeopyle of the paratype of imple tosphaeridium transfodum, the ‘type species’, is probably apical. but the morphology of the holotype remains uncertain. to quote fensome et al. (2009, p. 38): “this genus is useful, if of dubious status, since it serves as a repository for chorate forms whose general morphology accords with a gonyaulacacean dinoflagellate affinity, but whose archaeopyle is uncertain.” impletosphaeridium apodastum sp. nov. (plate 9, figs 5–8) holotype. plate 9, fig. 5, from a cuttings sample at 4565– 4575 m in hekja o-71, gsc type collection no. 137903, sample p18737, slide 01, co-ordinates 13.3 × 108.0, england finder n37/3. central body length 30 μm, width 26 μm; length of processes up to 17 μm, width less than 1 μm. the age determined for the sample from which the holotype was recovered is late danian. etymology. the epithet is from the greek apodastos, meaning separated or apportioned, in reference to the bifurcations of the processes distally. diagnosis. a species of impletosphaeridium with solid delicate processes that divide distally into two relatively long branches. the branches are bifurcate at the tips and usually curve back towards the central body. description. this species has a rounded, small central body that is granulate. distally, the solid delicate pro cesses divide into two, relatively long branches, each of which is bifurcate at its tip. these distal branches usually curve back towards the central body. the initial bifurcations are generally one-quarter to one-third as long as the main stem of the process but can be as much as a half or as little as a quarter the length. there are about 50 processes per specimen. size. central body length 20–33 μm, width 18–27 μm, length of processes 10–18 μm; six specimens measured. age. lo: selandian. remarks. impletosphaeridium apodastum is character ised by the high number of processes and their distinctive distal extremities, which branch. tips of the two branch es are usually bifid. occasional processes have three branches. because of the unknown nature of the archaeo py le, this species is included in impletosphaeridium. genus isabelidinium lentin & williams 1977b type. cookson & eisenack 1958, plate 4, fig. 10, as deflandrea korojonensis. 1976 isabelia lentin & williams, p. 56 (name illegitimate). 1977b isabelidinium lentin & williams, p. 167. 2009 isabelidinium lentin & williams – emend. fensome et al., p. 39. remarks. in recent decades, there has been much de bate about the definition of isabelidinium and similar genera such as manumiella and chatangiella, all of which have an isodeltaform, isothetaform and/or isoomegaform 2a plate whose partial or complete detachment forms the archaeopyle. species now included in chatangiella and isabelidinium can have an isodelata form, isothetaform or iso-omegaform 2a plate, but manu miella almost always has an isodeltaform 2a plate. much confusion remains regarding generic circumscriptions, but a resolution as to how these should be dealt with is beyond the scope of the present study. in our view, generic distinctions should be based mainly on variations in the shape and dimensions of the hexa 2a plate and secondarily on other expressions of tabulation such as the cingulum and on the ornamentation. chatangiella is clearly distinguished from isabelidi nium by its partite cingulum. fensome et al. (2009) emended manumiella, restricting it to forms with one symmetrically disposed antapical prominence. how ever, the antapex of the holotype of the type of the genus, bulletin36.qxp_bulletin 36 19/12/16 13.40 side 52 53 manumiella (originally broomea) seelandica (lange 1969, plate 3, fig. 3), is partly obscured so that its precise morphology cannot be determined; and the paratype (lange 1969, plate 2, fig. 10) clearly has two protuberances. thorn et al. (2009) emended the diagnosis of manumiella to allow for inclusion of a mes ophragm, which we would consider a non-diagnostic feature at generic rank. however, these authors did in dicate that manumiella was distinctive in being cir cumcavate, with isabelidinium being bicavate. for further comparisons, see under alterbidinium. isabelidinium cooksoniae (alberti 1959) lentin & williams 1977b 1959 deflandrea cooksoniae alberti, p. 97, plate 9, figs 1–6. 1961b isabelidinium belfastense cookson & eisenack, p. 71, plate 11, figs 4–6. 1967 australiella cooksoniae (alberti) – vozzhennikova, p. 132. 1976 isabelia cooksoniae (alberti) – lentin & williams, p. 57. 1976 isabelia belfastensis (cookson & eisenack) – lentin & williams, p. 57. 1977b isabelidinium cooksoniae (alberti) – lentin & williams, p. 167. 1977b isabelidinium belfastense (cookson & eisenack) – lentin & williams, p. 167. 1992 isabelidinium bujakii marheinecke, p. 86, 87, plate 18, figs 1–3; text-fig. 16. age. lo: maastrichtian. remarks. see fensome et al. (2009, p. 39) for a discussion of this species. isabelidinium cretaceum (cookson 1956) lentin & williams 1977b (plate 9, figs 15, 16) 1956 deflandrea cretacea cookson, p. 184, 185, plate 1, figs 1–4 (only). 1976 isabelia cretacea (cookson) – lentin & williams, p. 57. 1977b isabelidinium cretaceum (cookson) – lentin & williams, p. 167. 1983 manumiella? cretacea (cookson) – bujak & davies, p. 161. age. lo: maastrichtian. remarks. the holotype of this species (cookson 1956, plate 1, fig. 1 and helby et al. 1987, fig. 42l) and some of the other specimens illustrated in the protologue (cookson (1956, plate 1, figs 2–4) are all bicavate, with an endocyst that is broader than long. this bicavation, together with a general rounding of the apical and antapical regions, are considered here to be characteristic for isabelidinium cretaceum. however, the presence of a very short horn on two specimens (cookson 1956, plate 1, figs 3, 4), neither of which is the holotype, does not exclude their retention in this species. this species is retained in isabelidinium because it is bicavate (see discussion above under isabelidinium). isabelidinium microarmum (mcintyre 1975) lentin & williams 1977b (plate 9, fig. 17) 1975 deflandrea microarma mcintyre, p. 65, plate l, figs 5–8. 1976 isabelia microarma (mcintyre) – lentin & williams, p. 58. 1977b isabelidinium microarmum (mcintyre) – lentin & williams, p. 168. age. lo: early campanian. remarks. the holotype of isabelidinium microarmum (mcintyre 1975, plate 1, figs 5, 6) has a deltaform archae opyle as do all the other specimens (mcintyre 1975, plate 1, figs 6–8). genus kiokansium stover & evitt 1978 type. tasch et al. 1964, plate 3, fig. 8, as hystricho sphaeridium unituberculatum. 1978 kiokansium stover & evitt, p. 167. 1979 bacchidinium davey, p. 555. remarks. characteristic features of kiokansium are the precingular archaeopyle, which is formed from the loss of plates 3´´ and 4´´, and the dirigate to cauliflorate distal terminations of the numerous solid processes. kiokansium williamsii singh 1983 1983 kiokansium williamsii singh, p. 150, plate 54, figs 3–6. age. lo: cenomanian. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 53 54 genus kleithriasphaeridium davey 1974 type. davey 1974, plate 5, figs 1, 2, text-fig. 3, as klei thriasphaeridium corrugatum. 1974 kleithriasphaeridium davey, p. 55, 56. 1976 diversispinosa benson, p. 184. age. lo: coniacian. remarks. fensome et al. (2009, p. 40) emended the diagnosis of kleithriasphaeridium to include forms with mesotabular, tubular processes that are open or closed distally and “to circumscribe forms with a combination precingular–apical archaeopyle, a type of opening that occurs as an intraspecific variant of some species.” these authors also stressed that kleithriasphaeridium does not have fibrous processes, thus differentiating it from cordosphaeridium. kleithriasphaeridium mantellii (davey & williams 1966a) comb. nov. (plate 9, fig. 18) 1966a hystrichosphaeridium mantellii davey & williams, p. 66, plate 6, fig. 6. 1973 florentinia mantellii (davey & williams) – davey & verdier, p. 191. age. lo: coniacian. remarks. davey & williams (1966a, p. 66) stated that “the periphragm of processes [is] slightly fibrous.” on the same page they further noted that “… the peri phragm of the central body appears to be fairly heavily granular but on closer examination the granules apparently result from a fine reticulation” and that “an apical archaeopyle appears always to be present.” davey & verdier (1973, plate 4, figs 1, 3) re-illustrated the holotype of kleithrias phaeridium (as florentinia) mantellii and concluded that the archaeopyle is precingular, re sulting from the loss of the 3´´ plate. although the processes may be slightly fibrous, it is not obvious in the photographs. thus the characteristic feature of the species appears to be the or namentation of the main body. fensome & williams (2005, p. 48) noted that florentinia includes “forms with simple tubular processes such as florentinia aculeata and florentinia cooksoniae, as well as more ‘classic’ florentinia types, such as florentinia laciniata and florentinia ferox, with more complicated processes.” they recommended that taxa with simple tubular processes and precingular rather than combination archaeopyles, such as florent inia cooksoniae, should be transferred to kleithrias phaeri dium. thus the new com bination kleithriasphaeridium mantellii is proposed here. genus laciniadinium mcintyre 1975 type. mcintyre 1975, plate 4, figs 12, 13, as lacini a dinium orbiculatum. 1975 laciniadinium mcintyre, p. 70. 1984 sinocysta he chengquan, p. 769, 773. synopsis. a proximate peridiniacean (palaeoperidinioidean) cyst that is subspherical to biconical and com pressed dorso-laterally; when biconical, single horns are developed at the apical and antapical poles. acavate. autophragm smooth or ornamented with granules or echinae. cingulum indicated by low ridges. archaeo pyle combination intercalary-precingular, with formula ip(1–3a + 3–5´´), operculum simple, attached along the anterior margin of the cingulum. remarks. laciniadinium differs from palaeoperidinium in having one rather than two antapical horns, and an archaeopyle involving intercalary and precingular plates only. laciniadinium arcticum (manum & cookson 1964) lentin & williams 1980 (plate 9, figs 19, 20) 1964 diconodinium arcticum manum & cookson, p. 18, 19, plate 6, figs 1–4. 1980 laciniadinium arcticum (manum & cookson) – lentin & williams, p. 41. 1986 laciniadinium williamsii ioannides, p. 28, plate 10, figs 1–6; plate 11, fig. 5. age. lo: late maastrichtian. remarks. in their description of laciniadinium (as diconodinium) arcticum, manum & cookson (1964) stated that the autophragm was ornamented with minute granules up to 0.5 μm in diameter. the size range of the species was: length 50–73 μm, breadth 32– 53 μm. ioannides (1986) erected the species lacinia di nium williamsii, which can have a smooth to finely or namented autophragm; laciniadinium williamsii va ries from 39 to 54 μm in length and 31 to 43 μm in width. unfortunately, ioannides (1986) did not com pare laciniadinium williamsii to laciniadinium ar c ti bulletin36.qxp_bulletin 36 19/12/16 13.40 side 54 55 cum, although he did differentiate it from lacinia dinium orbiculatum mcintyre 1975 and laciniadinium biconiculum mcintyre 1975. the only difference that we can determine for separating laciniadinium arcti cum and laciniadinium williamsii is on the size, but this overlaps as demonstrated above. accordingly, we herein consider laciniadinium williamsii to be a taxonomic junior synonym of laciniadinium arcticum. genus lentinia bujak in bujak et al. 1980 type. bujak et al. 1980, plate 18, figs 7–9, text-fig. 18a–f, as lentinia serrata. 1980 lentinia bujak in bujak et al., p. 69. remarks. fensome et al. (2009, p. 42) provided a com prehensive synopsis for lentinia that includes reference to the large 2a isodeltaform archaeopyle, which dominates the dorsal epicystal area of the pericyst. lentinia serrata bujak in bujak et al. 1980 (plate 10, fig. 9) 1980 lentinia serrata bujak in bujak et al., p. 71, 72, plate 18, figs 7–12; text-figs 18a–f, 19. age. lo: latest priabonian. genus licracysta fensome et al. 2007 type. fensome et al. 2007, plate 4, figs 9–12, as licra cysta corymbus. 2007 licracysta fensome et al., p. 400, 402. remarks. licracysta is an areoligeracean cyst with non tabular to penitabular processes, many dolabrate to moderately licrate, that are absent from the dorsoventral region. in glaphyrocysta, the processes are not licrate or dolobrate distally. also, processes tend to be longer relative to the size of the central body than in licracysta. a synopsis for licracysta is provided by fen some et al. (2009, p. 42). licracysta corymbus fensome et al. 2007 (plate 10, fig. 16) 2007 licracysta corymbus fensome et al., p. 402, 404, 406, 408; plate 4, figs 5, 6, 9–20; plate 5, figs 1–8, 12, 16, 20. age. lo: rupelian. licracysta? semicirculata (morgenroth 1966b) fensome et al. 2007 1966b cyclonephelium semicirculatum morgenroth, p. 9, 10, plate 2, figs 3, 4. 1978 areoligera semicirculata (morgenroth) – stover & evitt, p. 18. 2007 licracysta? semicirculata (morgenroth) – fen some et al., p. 408. age. lo: latest rupelian. remarks. the distribution of the process complexes in licracysta? semicirculata is reminiscent of the late cretaceous forms of areoligera, in which the midventral surface lacks ornamentation, ventral processes being restricted to ambital linear complexes. genus lingulodinium wall 1967 type. deflandre & cookson 1955, plate 9, fig. 6, as hystrichosphaeridium machaerophorum. 1967 lingulodinium wall, p. 109. synopsis. a chorate gonyaulacacean (cribroperidinioid) cyst with a subspherical central body. acavate. the cen tral body is smooth or ornamented with features of low relief and bears numerous simple, hollow, apparently nontabular processes that are variable in length and distal ending. archaeopyle precingular, resulting from loss of a variable number of precingular plates (one to five), or epicystal. when the archaeopyle is precingular, the opercular pieces are usually free and separate. remarks. we do not know if the variation in archaeopyle development among specimens of lingulodinium bears any relationship to age or palaeoenvironment. lingulodinium funginum (morgenroth 1966a) islam 1983a (plate 10, figs 1, 2) 1966a baltisphaeridium funginum morgenroth, p. 17, 18, plate 3, figs 7, 8. 1983a lingulodinium funginum (morgenroth) – islam, p. 341. age. peak of lingulodinium spp. inconsistent within eocene – it probably has greater palaeoenvironmental than biostratigraphic significance. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 55 56 remarks. according to morgenroth (1966a), lingulo dinium funginum is characterised by having a variable number of processes that are distally bulbose, in contrast to species such as lingulodinium machaerophorum that have acuminate process endings. specimens of lingulodinium with bulbose process endings are com mon in some samples from labrador margin wells and these are included here in lingulodinium funginum. however, there is a caveat. kokinos & anderson (1995), in laboratory experiments, noted that the length and nature of the processes of resting cysts of the extant lingulodinium polyedrum show considerable variation. the processes of the cysts, more familiarly known to palynologists as lingulodinium machaerophorum, could be distally acuminate, bulbose, smooth or with small granules or spines; process length could vary up to a maximum of 10–12 μm. based on their findings, ko kinos & anderson (1995) considered lingulodinium funginum, lingulodinium sadoense and lingulodinium brevi spinosum to be taxonomic junior synonyms of lingulodinium machaerophorum. because morphological types included here in lingulodinium funginum occur only in the eocene in the study material, we propose to retain the species as defined in morgenroth (1966a) and prefer to avoid the use of infraspecific ranks (e.g. variety). retention may also provide useful information on pa laeo environments. lingulodinium machaerophorum (deflandre & cookson 1955) wall 1967 (plate 10, figs 3, 4) 1955 hystrichosphaeridium machaerophorum de flandre & cookson, p. 274, plate 9, figs 4, 8. 1961 baltisphaeridium machaerophorum (deflandre & cookson) – gerlach, p. 191, 192. 1966 cleistosphaeridium machaerophorum (de flandre & cookson) – davey et al., p. 170; combination not validly published. 1967 lingulodinium machaerophorum (deflandre & cookson) – wall, p. 109. age. lo not confirmed in the study area, but globally the species extends to the present day. genus nyktericysta bint 1986 type. bint 1986, plate 4, figs 1, 2, 5, 6, text-fig. 3a, b, as nyktericysta davisii. 1986 nyktericysta bint, p. 148, 149. 1986 balmula bint, p. 158. 1999 quantouendinium mao shaozhi et al., p. 155, 156. remarks. fensome et al. (2009, p. 46) emended the generic diagnosis of nyktericysta and considered the two wall layers to be the endophragm and ectophragm, to better facilitate comparison with closely similar genera. the ectophragm is finely perforate. according to the emendation, nyktericysta always has one apical and two antapical horns. in addition, one or two lateral equatorial horns commonly occur, and these may have preand postcingular extensions. the archaeopyle is apical, with the formula a(1–4´) and the operculum usually remains attached. fensome et al. (2009, p. 46) noted that vesperopsis differs from nyktericysta in not having an ectophragm, and that endoceratium and pseudo ceratium differ in having only an apical, postcingular and antapical horn rather than two antapical horns. mao shaozhi et al. (1999, p. 156) differentiated quantouendinium from nyktericysta bint 1986 and ves per opsis bint 1986 on the number and nature of the horns. fensome et al. (2009, p. 46) in their emendation of nyktericysta stated that it always has two antapical horns and commonly can have one or two lateral, equatorial horns. this indicates that the main distinction between quantouendinium and nyktericysta is that the former has one postcingular horn. however, from mao shaozhi et al. (1999, fig. 4), the location of both posterior horns appears to be antapical. we thus consider quantouendinium to be a taxonomic junior synonym of nyktericysta. nyktericysta davisii bint 1986 (plate 10, figs 5, 6) 1986 nyktericysta davisii bint, p. 149, 150, 152, 153, plate 4, figs 1–12; plate 8, figs 1–6; textfigs 3a–c, 4a, b, 10a, b. age. lo: late albian. remarks. nyktericysta davisii is characterised by the presence of five horns, of which the two lateral equatorial have preand postcingular extensions. the ectophragm is perforate. tabulation may be indicated by low sutural ridges. nyktericysta dictyophora he chengquan et al. 1992 (plate 10, figs 7, 8) bulletin36.qxp_bulletin 36 19/12/16 13.40 side 56 57 1992 nyktericysta dictyophora he chengquan et al., p. 184, 190, 191, plate 1, figs 1–9. 1992 nyktericysta dictyophora subsp. circularis he chengquan et al., p. 185, 191, plate 1, figs 7–9. 1992 nyktericysta fusiformis he chengquan et al., p. 185, 191, 192, plate 1, fig. 16; plate 2, figs 1–3. 1999 quantouendinium dictyophorum (he cheng quan et al.) – mao shaozhi et al., p. 156. age. lo: cenomanian. remarks. mao shaozhi et al. (1999, p. 157) considered nyktericysta dictyophora subsp. circularis to be a junior synonym of quantouendinium dictyophorum (that is, with the autonym). they also considered nyktericysta fusiformis to be taxonomic junior synonyms of this species. nyktericysta tripenta (bint 1986) fensome et al. 2009 (plate 10, figs 11, 12) 1986 balmula tripenta bint, p. 158, 160, plate 6, figs 9–17; plate 7, fig. 8; text-fig. 6a, b. 2009 nyktericysta tripenta (bint) – fensome et al., p. 46. age. lo: albian. remarks. fensome et al. (2009, p. 46) noted that nyk tericysta tripenta has a coarse autophragmal reticulum, with muri being up to 10 μm wide. genus odontochitina deflandre 1937 type. deflandre 1937, plate 18 (also labelled plate 15), fig. 8, as odontochitina silicorum. 1937 odontochitina deflandre, p. 94. remarks. the synopsis provided for odontochitina by fen some et al. (2009, p. 46) does not allow for the inclusion of forms with abbreviated horns. odontochitina ancala bint 1986 (plate 10, fig. 14) 1986 odontochitina ancala bint, p. 139, 140, plate 1, figs 2–8; plate 7, figs 1, 2; text-fig. 2a. age. lo: cenomanian. remarks. according to bint (1986, p. 140), “odonto chitina ancala differs from o. operculata by having an elbow and cingular notch in the right lateral horn, localised perforations about midway along the horns, and an elongate ventral extension of the antapical pericoel.” odontochitina costata alberti 1961 (plate 10, fig. 15) 1961 odontochitina costata alberti, p. 31, plate 6, figs 10–13. 1962 odontochitina striatoperforata cookson & eisenack, p. 490, plate 3, figs 14–19. age. lo: latest campanian. remarks. odontochitina costata was emended by clarke & verdier (1967, p. 58, 59), who considered odonto chitina striatoperforata to be intergradational with, and a taxonomic junior synonym of, odontochitina costata. odontochitina porifera cookson 1956 (plate 10, fig. 13) 1956 odontochitina porifera cookson, p. 188, plate 1, fig. 7. age. lo: santonian. remarks. in the scotian margin wells, the lo of odon tochitina porifera is also santonian (fensome et al. 2009, p. 47). genus oligosphaeridium davey & williams 1966a type. white 1842, plate 4, fig. 11, as xanthidium tubi ferum var. complex. 1966a oligosphaeridium davey & williams, p. 70, 71. remarks. in their synopsis for oligosphaeridium, fen some et al. (2009, p. 47) stated: “processes more or less equal in size and general shape.” there are two exceptions, the first and fourth apical (1´, 4´) are generally slender compared to the other apicals, although all the apicals tend to be slender. oligosphaeridium albertense (pocock 1962) davey & williams 1969 (plate 10, fig. 17) bulletin36.qxp_bulletin 36 19/12/16 13.40 side 57 58 1962 hystrichosphaeridium albertense pocock, p. 82, plate 15, figs 226, 227. 1962 hystrichosphaeridium irregulare pocock, p. 82, plate 15, figs 228, 229. 1964 hystrichosphaeridium coelenteratum tasch in tasch et al., p. 195, plate 2, fig. 11. 1964 hystrichosphaeridium dispare tasch in tasch et al., p. 195, plate 2, fig. 8 1964 hystrichosphaeridium reniforme tasch in tasch et al., p. 193, plate 2, fig. 6. 1969 oligosphaeridium albertense (pocock) – davey & williams, p. 5. age. lo: early cenomanian. remarks. in his emendation of oligosphaeridium alber tense, brideaux (1977, p. 27, 28) described the processes as hollow, flared to tubiform, with the open distal margins being “variably secate, occasionally aculeate or serrate; the distal third of some processes variably fene strate ….” jansonius (1986, p. 213) described the holotype of oligosphaeridium albertense, which he reillustrated (his plate 4, figs 4, 5), as having processes that were proximally nearly cylindrical but widened, and were strongly flared distally. the distal margins of the processes were “occasionally scalloped, carrying numerous coarse to fine, slender spinules.” jansonius (1986) considered hystricho s phaeridium (as oligosphaeridium) irregulare to be a taxonomic junior synonym of oligosphaeridium albertense. stover & evitt (1978, p. 68, 69) regarded hystricho sphae ridium (as oligosphaeridium) coelenteratum, hy stri chos phaeridium (as oligosphaeridium) dispare and hystrichosphaeridium (as oligosphaeridium) reniforme as taxonomic junior synonyms of oligosphaeridium irre gulare. thus, by implication, oligosphaeridium coelenteratum, oligosphaeridium dispare and oligosphaeri dium reniforme are all taxonomic junior synonyms of oligo sphaeridium albertense. oligosphaeridium pulcherrimum (deflandre & cookson 1955) davey & williams 1966a (plate 10, figs 18, 19) 1955 hystrichosphaeridium pulcherrimum deflandre & cookson, p. 270, 271, plate 1, fig. 8; textfigs 21, 22. 1966a oligosphaeridium pulcherrimum (deflandre & cookson) – davey & williams, p. 75, 76. age. lo: regional – santonian; lad – danian. not plotted. remarks. included here in oligosphaeridium pulcherrimum are those forms that have a mixture of perforated and unperforated processes. oligosphaeridium complex has no perforated processes. oligosphaeridium totum brideaux 1971 (plate 10, fig. 20) 1971 oligosphaeridium totum brideaux, p. 88, 89, plate 25, figs 53–55, 57. 1971 oligosphaeridium diastema singh, p. 337, plate 55, figs 4, 5; plate 56, figs 1, 2. age. lo: early cenomanian. remarks. based on nannofossil control, fensome et al. (2009) considered the last occurrence for oligo sphae r idium totum to be in the early cenomanian. genus operculodinium wall 1967 type. deflandre & cookson 1955, plate 8, figs 3, 4, as hystrichosphaeridium centrocarpum. 1967 operculodinium wall, p. 110, 111. remarks. operculodinium is a chorate gonyaulacacean (criboperidinioidean) cyst with a spheroidal to slightly ovoidal central body with a reticulate wall. the central body bears numerous nontabular to contabular processes. processes are generally solid, of the same size in individual species and distally bifid to aculeate. the archae opyle is precingular, with the formula p3´´ ; the operculum is free. fensome et al. (2009, p. 48) dis cussed the differences between operculodinium and the closely similar genus exochosphaeridium. operculodinium centrocarpum (deflandre & cookson 1955) wall 1967 (plate 11, figs 5, 6) 1955 hystrichosphaeridium centrocarpum deflandre & cookson, p. 272, 273, plate 8, figs 3, 4. 1961 baltisphaeridium centrocarpum (deflandre & cookson) – gerlach, p. 192,193. 1965 cordosphaeridium centrocarpum (deflandre & cookson) – de coninck, p. 33. 1966a cordosphaeridium tiara subsp. centrocarpum (deflandre & cookson) – morgenroth, p. 26. 1967 operculodinium centrocarpum (deflandre & cookson) – wall, p. 111. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 58 59 1969 cordosphaeridium? microtriainum subsp. centrocarpum (deflandre & cookson) – de coninck, p. 32. 1978 cleistosphaeridium centrocarpum (deflandre & cookson) – jiabo, p. 61. 1983 operculodinium echigoense matsuoka, p. 126, plate 7, figs 1, 2a, b, 3–5, 8. 1987 operculodinium? echigoense matsuoka – mudie, p. 804. age. lo: tortonian. genus palaeocystodinium alberti 1961 type. alberti 1961, plate 7, fig. 12, as palaeo cysto di nium golzowense. 1961 palaeocystodinium alberti, p. 20. 1963 cystodiniopsis vozzhennikova, p. 185. remarks. in their emendation of palaeocystodinium, fen some et al. (2009, p. 48) stated that they are “peri diniacean (deflandreoid) cysts that are fusiform in shape, with single prominent pointed horns apically and antapically; the horns are generally long and there may be a short accessory antapical horn.” fensome et al. (2009) differentiated palaeo cystodinium from the genus sval bardella manum 1960 on the nature of the horns distally; the latter has bluntly rounded apical and antapical horns. palaeocystodinium bulliforme ioannides 1986 (plate 11, figs 1, 2) 1986 palaeocystodinium bulliforme ioannides, p. 31, plate 17, figs 2–5. age. lo: selandian. palaeocystodinium golzowense alberti 1961 (plate 10, fig. 10) 1961 palaeocystodinium golzowense alberti, p. 20, plate 7, figs 10–12; plate 12, fig. 16. age. lo: late tortonian (late miocene). palaeocystodinium teespinosum fensome et al. 2009 (plate 11, fig. 3) 2009 palaeocystodinium teespinosum fensome et al., p. 50, plate 8, figs m–p. age. lo: early rupelian. remarks. palaeocystodinium teespinosum differs from pa laeo cystodinium golzowense in having delicate tshaped spinelets that are about 2 μm long and are especially common on the apical and antapical horns. genus palaeohystrichophora deflandre 1935 type. deflandre 1935, plate 8, fig. 4, as palaeo hy strichophora infusorioides. 1935 palaeohystrichophora deflandre, p. 230. remarks. in their remarks for this genus, fensome et al. (2009, p. 50) compared palaeohystrichophora to sub ti lisphaera jain & millepied 1973. both genera lack an obvious archaeopyle, but harker (1979, p. 374, fig. 1) observed a combination archaeopyle with the formula 3i3p(1–3a + 3–5´´) in palaeohystrichophora infusorioides; according to harker, the operculum remains attached along the posterior or cingular margin. bujak & davies (1983, p. 62, text-fig. 4) observed an archaeopyle, which they termed the transverse archaeopyle, in some species assigned to subtilisphaera. this archaeopyle resulted from the development of sutures between the apicals and three anterior intercalary plates and between the anterior lateral margins of the intercalaries. the major observable difference between the two genera is the presence of processes on the pericyst of palaeohystrichophora. forms with pro cesses that are presently included in subtilisphaera should probably be included in palaeohystrichophora. palaeohystrichophora infusorioides deflandre 1935 (plate 11, fig. 4) 1935 palaeohystrichophora infusorioides deflandre, p. 230, 231, plate 8, fig. 4. 1943 palaeohystrichophora paucisetosa deflandre, p. 507, 508, text-fig. 26. age. lo: campanian. genus palaeoperidinium deflandre 1934 ex sarjeant 1967 type. ehrenberg 1838, plate 1, fig. 4, as peridinium pyrophorum. 1934 palaeoperidinium deflandre, p. 968; name not validly published. 1963 pentagonum vozzhennikova, p. 183; name not validly published. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 59 60 1967 palaeoperidinium deflandre ex sarjeant, p. 246, 247. 1967 pentagonum vozzhennikova ex vozzhen nikova, p. 106; name illegitimate. 1970 astrocysta davey, p. 359. remarks. the synopses of palaeoperidinium provided by stover & evitt (1978, p. 217) and fensome et al. (2009, p. 51) are very similar. in their remarks, fen some et al. (2009) took into account the emendation of palaeoperidinium by evitt et al. (1998, p. 46, 48), who noted that the most commonly preserved wall of the cyst is an exophragm, which is unusual in having been formed outside but in contact with the exterior surface of the theca. palaeoperidinium pyrophorum (ehrenberg 1838 ex wetzel 1933b) sarjeant 1967 (plate 11, figs 7, 8) 1838 peridinium pyrophorum ehrenberg, plate 1, figs 1, 4; name not validly published. 1933b peridinium pyrophorum ehrenberg ex wetzel, p. 164, 165. 1967 palaeoperidinium pyrophorum (ehrenberg ex wetzel) – sarjeant, p. 246. 1967 peridinium basilium drugg, p. 13, plate 1, figs 9–11; plate 9, fig. 1a, b. 1967 pentagonum marginatum vozzhennikova, p. 107, plate 46, figs 1, 3, 4, 6; generic name illegitimate. 1967 pentagonum sibiricum vozzhennikova, p. 106, 107, plate 46, figs 2, 5. 1967 peridinium conicum var. larjakiense vozzhen nikova, p. 71, 72, plate 16, figs 1a, b, 2a, b. 1973 palaeoperidinium deflandrei lentin & wil liams, p. 105. 1981 palaeoperidinium larjakiense (vozzhennikova) – lentin & williams, p. 210. age. lo: selandian; peak: early danian. remarks. gocht & netzel (1976, p. 403–405) and evitt et al. (1998, p. 48, 49) provided comprehensive, concise emendations of palaeoperidinium pyrophorum, and re solved its relationship to the thecate equivalent. genus palynodinium gocht 1970 type. gocht 1970, fig. 4, nos 1a–c, as palynodinium grallator. 1970 palynodinium gocht, p. 135, 137, 138, 140. palynodinium grallator gocht 1970 (plate 11, figs 9, 10) 1970 palynodinium grallator gocht, p. 135, 137, 138, 140, fig. 2a–e; fig. 4, nos 1a–c, 2a, b, 3a, b, 4a, b, 5a, b, 6a, b, 7, 8; fig. 5, nos 1, 2a, b. age. lo: latest maastrichtian. genus petalodinium williams et al. 2015 type. williams & downie 1966b, plate 20, figs 1, 2, as wetzeliella condylos. 2015 petalodinium williams et al., p. 307. remarks. williams et al. (2015) erected the genus peta lo dinium for wetzelielloidean dinocysts with a latiepeliform archaeopyle and a pericyst that is smooth or or namented with features of low relief. petalodinium condylos (williams & downie 1966b) williams et al. 2015 (plate 11, figs 11, 12) 1966b wetzeliella condylos williams & downie, p. 193, 194, plate 20, figs 1, 2. 1976 rhombodinium condylos (williams & downie) – lentin & williams, p. 128. 1979 dracodinium condylos (williams & downie) – costa & downie, p. 43. 2015 petalodinium condylos (williams & downie) – williams et al., p. 308. age. lo: ypresian. remarks. as williams et al. (2015) have demonstrated, the holotype of petalodinium condylos (williams & dow nie 1966b, plate 20, fig. 1, as wetzeliella condylos), which is the type of petalodinium, has a latiepeliform archaeopyle. apparently related to the latiepeliform archaeopyle is the nature of the apical horn, which in petalodinium condylos is reduced or absent, the apex being rounded. the pericyst is verrucate to tuberculate. often, the intratabular ornamentation delineates the tabulation: other verrucae or tubercles are penitabular. both periand endophragm are thick, about 3 μm. some specimens observed in the labrador margin samples are devoid of ornamentation, but are otherwise identical. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 60 61 genus phelodinium stover & evitt 1978 type. corradini 1973, plate 28, fig. 3, as deflandrea pentagonalis. 1978 phelodinium stover & evitt, p. 117, 118. synopsis. a dorso-ventrally compressed, proximate protoperi diniacean (protoperidinioidean) cyst with one apical and two antapical horns. the cyst is pentagonal with convex to concave lateral sides. it is produced into one apical and two antapical horns that can be of variable length, are usually acuminate distally, and may have solid tips. cornucavate where two wall layers can be discerned. cingulum sometimes present. endocyst, where observable, and pericyst smooth or ornamented with features of low relief. archaeopyle intercalary, with formula i(2a), the second anterior intercalary plate is deltaform; operculum usually free. remarks. the emendation of phelodinium by mao shaozhi & norris (1988, p. 51, 52) related to the archaeopyle index, given in the original diagnosis as 0.3 to 0.4 by stover & evitt (1978, p. 117). mao shaozhi & norris (1988) expanded the index to between 0.3 and 0.6; these authors also classified the archaeopyle as standard hexa 2a. one difficulty with phelodinium is determining if it has one or two wall layers. according to stover & evitt (1978, p. 118) “phelodinium differs from lejeunia in being cavate and in having a peridinioid outline with straight to concave sides. lejeunia has an autophragm only, and its lateral margins are normally convex.” that there are two wall layers in phelodi nium is often extremely difficult to discern and the outline of the cyst in the two genera appears to be variable. phelodinium kozlowskii (górka 1963) lindgren 1984 (plate 12, figs 1–3) 1963 lejeunia kozlowskii górka, p. 41, plate 5, fig. 4. 1970 astrocysta kozlowskii (górka) – davey, p. 369. 1977 senegalinium kozlowskii (górka) – harland, p. 189. 1984 phelodinium kozlowskii (górka) – lindgren, p. 181. age. lo: danian. remarks. harland (1973, p. 673) and harker & sarjeant (1975, p. 223) considered phelodinium kozlowskii to be a taxonomic junior synonym of phe lodinium tricuspe. in their emendation of phelodinium tricuspe, lejeune-car pentier & sarjeant (1981, p. 20) stated that the “phragma [is] apparently composed of a single layer (autophragm).” these authors also ob served that in the holotype of phelodinium (as leje unecysta) tricuspe only one wall layer could be discerned: this observation explains why these authors accepted the transfer of the species to lejeune cysta by artzner & dörhöfer (1978). lindgren (1984, p. 181–183) recorded two wall layers in phelo dinium kozlowskii, with the endophragm being closely appressed to the periphragm except at the tips of the horns. some of the specimens of phelodinium kozlowskii encountered in this study show a similar morphology, which is why the species is retained in phelodinium and separated from phelodinium tricuspe; assuming that the holotype of the latter species has a single wall layer, it should be included in lejeunecysta. genus phthanoperidinium drugg & loeblich jr. 1967 type. drugg & loeblich jr. 1967, plate 1, fig. 4, as phtha noperidinium amoenum. 1967 phthanoperidinium drugg & loeblich jr., p. 182. remarks. fensome et al. (2009, p. 54) provided a com prehensive synopsis of phthanoperidinium, which in cluded the findings by edwards & bebout (1981, p. 36) and islam (1982, p. 306) on the variability and complexity of the archaeopyle. although the variability has not been noted rigorously, if the trend is similar to that in other peridiniaceans, it is probable that the various types have different stratigraphic ranges. this has already been demonstrated in phthanoperidinium geminatum and phthanoperidinium regale, whose only dif ference is in the nature of the archaeopyle. phthanoperidinium coreoides (benedek 1972) lentin & williams 1976 (plate 11, figs 13, 14) 1972 hystrichogonyaulax coreoides benedek, p. 20, plate 9, fig. 4a–c. 1976 phthanoperidinium coreoides (benedek) – lentin & williams, p. 76. 1976 phthanoperidinium tritonium eaton, p. 299, 300, plate 17, figs 2, 3, 6, 7; text-figs 23c, 24a, b. age. lo: rupelian. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 61 62 remarks. following fensome et al. (2009, p. 55), phtha no peridinium coreoides and phthanoperidinium comatum are separate species, distinguished on the relative length of the processes. in phthanoperidinium coreoides, the pro cesses are about one third the diameter of the central body; in phthanoperidinium comatum, the processes are about one half the diameter of the central body. phthanoperidinium levimurum bujak in bujak et al. 1980 (plate 11, fig. 15) 1980 phthanoperidinium levimurum bujak in bujak et al., p. 74, plate 19, figs 13–16; text-figs 20e, 22b. age. lo: priabonian. remarks. phthanoperidinium levimurum is characterised by having the tabulation delineated by membranes, which are usually smooth distally but may sometimes be denticulate. phthanoperidinium multispinum bujak in bujak et al. 1980 (plate 11, figs 16, 17) 1980 phthanoperidinium multispinum bujak in bujak et al., p. 74, plate 19, figs 17–19; textfig. 20f. age. lo: latest priabonian. phthanoperidinium regale bujak 1994 (plate 11, figs 18, 19) 1994 phthanoperidinium regale bujak, p. 130, plate 4, figs 4–6. age. lo: lutetian. remarks. phthanoperidinium regale has ornamentation similar to phthanoperidinium geminatum bujak in bujak et al. 1980, but differs in having a combination archae opyle, with the formula ip(2a+4´´). phano per idinium geminatum has the more usual archaeopyle type for the genus, losing the 2a plate only. in phthanoperidinium regale, the operculum is free. phthanoperidinium stockmansii (de coninck 1975) lentin & williams 1977a (plate 11, fig. 20) 1975 peridinium stockmansii de coninck, p. 97, 98, plate 17, figs 18–37. 1977a phthanoperidinium stockmansii (de coninck) – lentin & williams, p. 131. 1976 phthanoperidinium echinatum eaton, p. 298, 299, plate 17, figs 8, 9, 12; text-fig. 23b. 1980 phthanoperidinium? pseudoechinatum bujak in bujak et al., p. 75, 76, plate 19, fig. 20; textfig. 20c. age. lo: priabonian. remarks. phthanoperidinium stockmansii has penitabular rows of small processes that distally are clavate. rows of processes usually occur in parallel pairs, somewhat like railway tracks. de coninck (1977, p. 40) considered phtha noperidinium echinatum to be a junior taxonomic synonym of phthanoperidinium stockmansii. islam (1982, p. 315) agreed with this synonymy, but mistakenly thought that phthanoperidinium echinatum was the senior name. accepting islam’s synonymy of phthano peri dinium pseudoechinatum with phthanoperi dinium echinatum, the former thus becomes a junior taxonomic synonym of phtha no peridinium stockmansii. genus piladinium williams et al. 2015 type. michoux 1988, plate 1, figs 2, 3, as kisselovia columna. 2015 piladinium williams et al., p. 308, 309. remarks. piladinium is a wetzelielloidean genus with a latiepeliform archaeopyle and processes connected by an ectophragmal membrane. for comparison with other wetzelielloidean genera, see williams et al. (2015). piladinium columna (michoux 1988) williams et al. 2015 (plate 12, figs 5, 6) 1988 kisselevia columna michoux, p. 28, 30, plate 1, figs 2, 3, 5, 6; plate 2, figs 3–5; text-fig. 7a, b. 1989 charlesdowniea columna (michoux) – lentin & vozzhennikova, p. 74. 2015 piladinium columna (michoux) – williams et al., p. 309. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 62 63 age. lo: ypresian. remarks. several of the observed labrador margin specimens of piladinium columna have reduced num bers of processes on the mid-ventral and mid-dorsal regions. the morphology of these specimens is thus approaching that of piladinium edwardsii. piladinium edwardsii (wilson 1967b) williams et al. 2015 (plate 12, figs 7, 8) 1967b wetzeliella edwardsii wilson, p. 477, figs 8, 9. 1978 kisselevia edwardsii (wilson) – stover & evitt, p. 111. 1989 charlesdowniea edwardsii (wilson) – lentin & vozzhennikovia, p. 227. 2015 piladinium edwardsii (wilson) – williams et al., p. 309. age. lo: early ypresian. not plotted. genus pseudoceratium gocht 1957 type. gocht 1957, plate 18, fig. 1, as pseudoceratium pelliferum. 1957 pseudoceratium gocht, p. 166. 1962 eopseudoceratium neale & sarjeant, p. 446. 1966a doidyx sarjeant, p. 205. synopsis. a dorso-ventrally compressed, proximate ceratiacean cyst with single apical, postcingular and antapical horns. wall formed of one or two layers. if two-layered, there can be endophragm and periphragm or endophragm and ectophragm. tabulation often indicated by ornamentation if present. ornamentation may be absent or it may be granular or consist of short processes, which may be trabeculate. cingulum sometimes obvious. archaeopyle apical, with formula a(1–4´), operculum free; sulcal notch offset to the left. pseudoceratium sp. (plate 12, fig. 4) remarks. our specimens of pseudoceratium differ from pseudoceratium pelliferum in having longer processes and in generally lacking a lateral horn. genus raphidodinium deflandre 1936 type. deflandre 1936, plate 10, figs 1, 2, 7, raphido dinium fucatum. 1936 raphidodinium deflandre, p. 184, 185. synopsis. chorate cyst with ovoidal central body and a clearly defined cingulum. the cingulum can subdivide the central body into equal epiand hypocysts or be anteriorly located, dividing the cyst into a short epicyst and much longer hypocyst. tabulation marked by mem branes or ridges and long, slender processes that are clustered along the cingulum and to a lesser extent in the antapical area. processes are distally blunt to multifurcate. archaeopyle indeterminate. remarks. marheinecke (1992, p. 79) considered the archaeopyle of raphidodinium fucatum subsp. com pac tum marheinecke 1992 to be possibly precingular. raphidodinium fucatum deflandre 1936 (plate 12, figs 9, 10) 1936 raphidodinium fucatum deflandre, p. 185, 186, plate 10, figs 1–7. age. lo: campanian. remarks. the unequal sizes of the epiand hypocysts suggest that raphidodinium fucatum could be a clado pyxiacean. genus reticulosphaera matsuoka 1983 type. matsuoka 1983, plate 4, fig. 8, as reticulato sphaera stellata. 1983 reticulatosphaera matsuoka, p. 116. reticulatosphaera actinocoronata (benedek 1972) bujak & matsuoka 1986 (plate 12, figs 11, 12) 1972 cleistosphaeridium actinocoronatum benedek, p. 34, plate 12, fig. 13; text-fig. 12. 1978 areosphaeridium? actinocoronatum (benedek) – stover & evitt, p. 20. 1983 reticulatosphaera stellata matsuoka, p. 116, 117, plate 4, figs 8–11; text-fig. 10. 1986 reticulatosphaera actinocoronata (benedek 1972) – emend. bujak & matsuoka 1986, p. 238. age. lo: earliest zanclian. not plotted. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 63 64 genus rhombodinium gocht 1955 emend. williams, damassa, fensome & guerstein in fensome et al. 2009 type. gocht 1955, text-fig. 1c, as rhombodinium draco. 1955 rhombodinium gocht, p. 85. 1961 wetzeliella subgenus rhombodinium (gocht) – alberti, p. 9. 2009 rhombodinium gocht 1955 emend. wil liams, damassa, fensome & guerstein in fensome et al., p. 57. remarks. we fully concur with the emendation in fensome et al. (2009), which emphasises the generic significance of the archaeopyle. according to the emended diagnosis, species of rhombodinium must have a solei form archaeopyle, which always has an operculum that is attached apically. rhombodinium was retained at generic rank by lentin & williams (1977a, p. 139). rhombodinium draco gocht 1955 (plate 12, figs 13, 14) 1955 rhombodinium draco gocht, p. 86, text-fig. 1a–c. 1961 wetzeliella subgenus rhombodinium draco (gocht) – alberti, p. 8. age. lo: priabonian. remarks. vozzhennikova (1967, p. 168) retained this taxon in rhombodinium. rhombodinium porosum bujak 1979 1979 rhombodinium porosum bujak, p. 314, 315, plate 1, figs 3, 5–8; plate 2, fig. 11; text-fig. 8c. age. lo: bartonian. genus rottnestia cookson & eisenack 1961a type. eisenack 1954, plate 9, fig. 5, as hystrichosphaera borussica. 1961a rottnestia cookson & eisenack, p. 40, 42. rottnestia borussica (eisenack 1954) cookson & eisenack 1961a (plate 13, fig. 1) 1954 hystrichosphaera borussica eisenack, p. 62, plate 9, figs 5a, b, 6, 7. 1961a rottnestia borussica (eisenack) – cookson & eisenack, p. 42. 1966a triblastula borussica (eisenack) – morgenroth, p. 15, 16. 1966b hystrichosphaeropsis borussica (eisenack) – sarjeant, p. 139; combination not validly published. age. lo: lutetian. not plotted. genus scalenodinium gen. nov. type. plate 12, fig. 15, as scalenodinium scalatum. etymology. the name is from the latin scalenus, meaning unequal, uneven, odd, in reference to the unequal variation in apical and antapical development. description. pericoel elongate, with an apical horn that is acuminate to rounded distally and generally a rounded antapex, although the latter may be extended into a short protuberance. when the endocyst is present, the cyst is bicavate. cingulum present or faintly expressed. periarchaeopyle intercalary, generally with formula i2a, and with a free operculum. the hexa 2a plate is steno deltaform. in some specimens, the apparent loss of the apical area (plate 12, figs 16–19) or of other intercalary plates besides the 2a (plate 12, fig. 15) indicates a com pound and/or combination archaeopyle. remarks. scalenodinium differs from isabelidinium in having a stenodeltaform, rather than a latito iso-ome ga form 2a plate. palaeocystodinium, which like scaleno dinium has a stenodeltaform 2a plate, differs from the latter genus in always having more or less apical and antapical horns and a periarchaeopyle that is always formed from the loss of a single plate, the 2a plate, rather than several plates as in scalenodinium. scalenodinium scalenum sp. nov. (plate 12, figs 15–20) holotype. plate 12, fig. 15, from a cuttings sample at 2135–2145 m in gilbert f-53, gsc type collection no. 137916, sample p39484, slide 01, co-ordinates 19.3 × 106.5, england finder t37/4. pericyst length 81 μm, width 31 μm. the age determined for the sample from which the holotype was recovered is ypresian. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 64 65 etymology. the name is from the latin scalenus, meaning unequal, uneven, odd, in reference to the unequal variation in apical and antapical development. description. a species of scalenodinium with an elongate pericoel that has an apex bearing a well-developed, distally acuminate to rounded apical horn and an antapex that is generally rounded, though an antapical protuberance may occasionally be developed. when an endocyst is present, the cyst is bicavate. the pericyst is verrucate to granulate in the mid-dorsal and midventral regions. a cingulum is rarely developed. both pericyst and endocyst have walls that are, at the most, 1 μm thick. periarchaeopyle intercalary, generally with formula i2a, operculum free. the hexa 2a plate is stenodeltaform. in some specimens, the apparent loss of the apical area (plate 12, figs 16–19) or of other intercalary plates besides the 2a (plate 12, fig. 15) indicates a compound and/or combination archaeopyle. size. pericyst length 77–86 μm, pericyst width 38–56 μm, apical horn length 18–27 μm, apical horn maximum breadth 11–14 μm. age. lo: ypresian. remarks. scalenodinium scalenum shows some variation in the antapical region: although generally rounded ant apically, some specimens have a prominent antapical horn that is more or less centrally located. the pericyst invariably has folds running across its surface. it is difficult to be specific about the nature of the archaeopyle since there is a common tendency for the apical polar area to break up. thus this breakup could be interpreted to denote a compound or combination archaeopyle. width and shape of the distal terminations of the horn or horns vary considerably. genus schematophora deflandre & cookson 1955 type. deflandre & cookson 1955, plate 6, figs 11, 12, as schematophora speciosa. 1954 schematophora deflandre & cookson, p. 1237 (name not validly published). 1955 schematophora deflandre & cookson, p. 262. remarks. the synopsis provided by fensome et al. (2009) covers all the salient points regarding the mor phology of schematophora. schematophora speciosa deflandre & cookson 1955 (plate 13, fig. 2) 1955 schematophora speciosa deflandre & cookson, p. 262, 263, plate 6, figs 11–13; plate 7, fig. 11. age. lo: priabonian. genus senegalinium jain & millepied 1973 type. jain & millepied 1973, plate 1, figs 1–3, as senegalinium bicavatum. 1973 senegalinium jain & millepied, p. 22, 23. synopsis. peridiniacean (deflandreoid) cysts with a peridinioid, usually elongate pericyst and two antapical horns that are more or less of equal length. bicavate with a circular to pentagonal endocyst. pericyst surface smooth or with low ornament, sometimes with cingulum indicated. periarchaeopyle intercalary, resulting from the loss of the isoto stenodeltaform deltaform 2a plate; the operculum commonly remains attached along the posterior margin. remarks. alterbidinium differs from senegalinium in having two antapical horns that are of unequal length. since alterbidinium contains species that may have a combination intercalary–precingular archaeopyle resulting from the loss of the 2a and 4´´ plate, it would not be surprising to find the same variability in senegalinium. senegalinium iterlaaense nøhr-hansen & heilmann-clausen 2001 (plate 13, figs 3, 4) 2001 senegalinium iterlaaense nøhr-hansen & heilmann-clausen, p. 164, 166–168, fig. 6, nos 1–6. age. lo: selandian. remarks. senegalinium iterlaaense differs from isabelidi nium viborgense in having a striate periphragm. genus senoniasphaera clarke & verdier 1967 type. clarke & verdier 1967, plate 14, fig. 8, as seno niasphaera protrusa. 1967 senoniasphaera clarke & verdier, p. 61. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 65 66 remarks. the cavate cysts included in senoniasphaera have the typical offset sulcal notch of areoligeraceans. the operculum of the apical archaeopyle is always free. senoniasphaera inornata (drugg 1970) stover & evitt 1978 (plate 13, fig. 5) 1970 chiropteridium inornatum drugg, p. 811, 812, fig. 3c–f. 1978 senoniasphaera inornata (drugg) – stover & evitt, p. 80. age. lo: danian. remarks. brinkhuis & schiøler (1996) recorded seno nias phaera inornata from the early danian of the geul hemmerberg cretaceous–palaeogene boundary section in limburg, south-eastern netherlands. this observation accords with the age of this species in sections at stevns klint (denmark) and alabama as recorded by hansen et al. (1986) and habib (1994) respectively. senoniasphaera inornata therefore appears to be a good danian index species. williams et al. (2004) placed the lo of the species close to the top of the danian. senoniasphaera microreticulata brideaux & mcintyre 1975 (plate 13, fig. 6) 1975 senoniasphaera microreticulata brideaux & mcintyre, p. 35, plate 11, figs 7–12; plate 12, figs 1–8. 1981 canningia microreticulata (brideaux & mcintyre) – below, p. 31. age. lo: cenomanian. remarks. lentin & williams (1981, p. 33) retained this species in senoniasphaera. senoniasphaera rotundata clarke & verdier 1967 (plate 13, fig. 7) 1967 senoniasphaera rotundata clarke & verdier, p. 62, 63, plate 14, figs 1–3; text-fig. 25. age. lo: latest campanian. remarks. in senoniasphaera rotundata, the endocyst does not protrude into the antapical horns. genus simplicidinium gen. nov. type. eaton 1976, plate 21, fig. 5, as impletosphae ri dium insolitum. etymology. the name derives from the latin simplicis, meaning simplicity, in reference to the relatively simple morphology of the cyst, comprising a spiny ball with an apical archaeopyle. diagnosis. proximochorate to chorate dinoflagellate cysts with a more or less symmetrical spheroidal to ovoidal central body. spines or processes numerous, isolated and non-tabulate, distally closed, with symmetrical distal terminations. archaeopyle apical, type 4a1´– 4´, operculum attached or detached. remarks. there seems to be considerable confusion currently surrounding the generic assignment of chorate species with non-tabulate spines/processes and a cryptic, or not clearly discernible or consistent archaeopyle. forms with a consistent apical archaeopyle and at least some asymmetrical processes are assignable to cleisto sphaeridium (eaton et al. 2001). islam (1993) proposed downiesphaeridium, purportedly for forms with an apical archaeopyle and simple non-tabulate processes. the processes of the type of downies phae ridium are typical of lin gulodinium, and the archaeopyle in the type of dow nies phae ridium looks precingular; it is thus suggested here that downies phae ridium may be a taxonomic junior synonym of lingu lo dinium, but at least should not be used beyond the type. many authors have assigned species that they consider to be chorate dinoflagellates but show no evidence of an archaeopyle to impleto sphaeridium, the type of that genus being suitably cryptic in its archaeopyle type. similar species deemed not to be dinoflagellates are as sign able to the acritarch genus balti sphaeridium. simplicidinium insolitum (eaton 1976) comb. nov. (plate 13, figs 9, 10) 1976 impletosphaeridium insolitum eaton, p. 308, plate 21, figs 5, 8; text-fig. 25b. 1978 cleistosphaeridium? insolitum (eaton) – stover & evitt, p. 31. age. local peak: earliest bartonian. remarks. eaton (1976, p. 308) commented on the presence of a polygonal opening in some of the specimens of simplicidinium (as impletosphaeridium) insolitum; such bulletin36.qxp_bulletin 36 19/12/16 13.40 side 66 67 openings are readily apparent in his ac companying illu strations (eaton 1976, plate 21, figs 5, 8). we thus assign this species, as type, to simplicidinium on the basis of the morphology of the holotype and because several of the specimens in this study also have an apical archaeopyle. genus sophismatia williams et al. 2015 type. williams & downie 1966a, plate 20, figs 2, 4, text-fig. 49, as wetzeliella tenuivirgula. 2015 sophismatia williams et al., p. 312, 313. remarks. williams et al. (2015) are followed here in re stricting sophismatia to wetzelielloideans with an equiepeliform 2a archaeopyle and trabeculate processes. sophismatia tenuivirgula (williams & downie 1966b) williams et al. 2015 (plate 13, fig. 8) 1966b wetzeliella tenuivirgula williams & downie, p. 188, 189, plate 19, figs 1, 2, 4, 5, 7; textfigs 49, 50. 1976 kisselevia tenuivirgula (williams & downie) – lentin & williams, p. 136. 1989 charlesdowniea tenuivirgula (williams & downie) – lentin & vozzhennikova, p. 227. 2015 sophismatia tenuivirgula (williams & downie) – williams et al., p. 313. age. lo: lutetian. remarks. the holotype of sophismatia tenuivirgula (wil liams & downie 1966b, plate 19, figs 2–4; text-fig. 49) clearly shows an equiepeliform archaeopyle, with the endoand periarchaeopyle being of almost identical size. the only difference is anteriorly, where the periarchaeopyle extends slightly beyond the endoarchaeopyle. genus spinidinium cookson & eisenack 1962 type. cookson & eisenack 1962, plate 1, figs 1, 2, as spinidinium styloniferum. 1962 spinidinium cookson & eisenack, p. 489. 2003 magallanesium quattrocchio & sarjeant, p. 138, 140. 2003 volkheimeridium quattrocchio & sarjeant, p. 136, 138. remarks. in their extensive review of spinidinium, voz z hennikovia and related genera, sluijs et al. (2009) emended the diagnosis of spinidinium. they restricted the genus to peridiniacean taxa with proximosutural spines, a stenoto isodeltaform 2a plate and an i2a archaeopyle, with the operculum typically attached posteriorly. spinidinium echinoideum (cookson & eisenack 1960a) lentin & williams 1976 (plate 13, figs 11, 12) 1960a deflandrea echinoidea cookson & eisenack, p. 2, plate 1, figs 5, 6. 1976 spinidinium echinoideum (cookson & eisenack) – lentin & williams, p. 64. 1978 vozzhennikovia echinoideum (cookson & eisenack) – stover & evitt, p. 130. 2003 spinidinium? echinoideum (cookson & eisenack) – quattrocchio & sarjeant, p. 136. age. lo: selandian. remarks. sluijs et al. (2009, p. 47) noted: “both the holo type (cookson & eisenack 1960a, plate 1, fig. 5) and specimens illustrated in sverdlove & habib (1974, plate 1, figs 3, 5, 6; text–fig. 1) show the 2a plate is iso deltaform to isothetaform. the 2a also remains at tached to the 4´´ plate. further the ornamentation is predominantly proximosutural. for the above reasons we include this species in spinidinium without question.” genus spiniferites mantell 1850 type. ehrenberg 1838, plate 1, fig. 5, as xanthidium ramosum, designated by davey & williams (1966b, p. 32), as lectotype of hystrichosphaera ramosa. 1850 spiniferites mantell, p. 191. 1933b hystrichosphaera wetzel, p. 33; name not validly published. 1937 hystrichosphaera wetzel ex deflandre, p. 61. 1953 hystrichokibotium klumpp, p. 387. remarks. spiniferites is characterised by having its gony aula coidean tabulation expressed by sutural ridges or membranes as well as gonal and sometimes sutural processes. distally, the gonal processes are always trifurcate, and the sutural processes are always bifurcate. the archaeopyle is precingular, with formula p3´´, with a free operculum. spiniferites ovatus matsuoka 1983 (plate 13, fig. 13) bulletin36.qxp_bulletin 36 19/12/16 13.40 side 67 68 1983 spiniferites ovatus matsuoka, p. 134, 135, plate 3, figs 1a–c, 2, 3a, b, 4a, b; text-fig. 19a, b. non 1984 spiniferites ovatus bujak, p. 192, plate 3, figs 15–18; illegitimate junior homonym. age. lo: messinian. remarks. spiniferites ovatus differs from spiniferites pseudofurcatus in having much shorter processes with shorter terminations. spiniferites pseudofurcatus (klumpp 1953) sarjeant 1970 (plate 13, fig. 14) 1953 hystrichokibotium pseudofurcatum klumpp, p. 388, plate 16, figs 12, 14. 1960 hystrichosphaera tertiaria eisenack & gocht, p. 515; text-fig. 4. 1966b hystrichosphaera buccina davey & williams, p. 42, 43, plate 4, fig. 1; text-figs 10, 11. 1969 hystrichosphaera pseudofurcata (klumpp) – gocht, p. 32. 1970 spiniferites pseudofurcatus (klumpp) – sarjeant, p. 76. age. lo in present study: late serravalian. lo according to piasecki (2003): tortonian. remarks. spiniferites pseudofurcatus differs from acho mo s phaera alcicornu only in having the tabulation clearly ex pressed by ridges. although so similar, the stratigraphic ranges of the two species differ considerably, with spini ferites pseudofurcatus having a much younger lo. spiniferites scabrosus (clarke & verdier 1967) lentin & williams 1975 (plate 13, fig. 15) 1967 hystrichosphaera scabrosus clarke & verdier, p. 49, 50, plate 9, figs 7–10; text-fig. 21. 1975 spiniferites scabrosus (clarke & verdier) – lentin & williams, p. 2155. age. lo: maastrichtian. remarks. the central body of spiniferites scabrosus has a distinctive scabrate to granulate wall and bears long, slender processes. in some specimens, the sutural ridges or membranes are not obvious, but such forms are retained here in spiniferites scabrosus. genus spongodinium deflandre 1936 type. ehrenberg 1838, plate 1, figs 1, 6, as peridinium delitiense. 1936 spongodinium deflandre, p. 169, 170. synopsis. gonyaulacacean (cribroperidinioid) cysts that are proximate, spheroidal to subpolyhedral, with a round ed antapex and usually an apical protrusion or horn. acavate. the wall is complexly reticulate to vesi culate, and its total thickness is sometimes greater along the cingulum and at the poles. wall atabulate or with structure/ornament ar ranged to reflect or suggest tabulation. archaeopyle precingular, with formula p3´´, oper culum free. remarks. the synopsis is similar to that provided by fensome et al. (2009, p. 60). an ectophragm is not developed but may be partially simulated on the cyst outline by the outer margins of membranes that constitute the wall. the genus samlandia is very similar in appearance to spongodinium, but the precise morphology of the type is not clear from available illustrations (eisenack 1954, plate 11, figs 12‒15). however, an sem micrograph of a specimen identified as samlandia chlamydophora by one of the authors shows that an occasionally perforate ectophragm connects processes in all but the sutural areas. such a morphology supports our interpretation that samlandia can thus be differentiated from spongodinium. spongodinium delitiense (ehrenberg 1838) deflandre 1936 (plate 13, figs 16, 20) 1838 peridinium delitiense ehrenberg, p. 110, plate 1, figs 1, 6. 1936 spongodinium delitiense (ehrenberg) – deflandre, p. 170, 171. age. lo: early danian. remarks. the labrador margin specimens of spongo di nium delitiense show considerable variation in size. the larger specimens have an lo in the maastrichtian and the smaller ones range up into the danian. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 68 69 spongodinium grossum (manum & cookson 1964) comb. nov. (plate 14, figs 1–4) 1964 chlamydophorella grossa manum & cookson, p. 17, 18, plate 5, figs 1, 2. 1986 chlamydophorella? grossa ioannides, p. 16. age. lo: campanian. remarks. in a description of spongodinium (as chlamy do pho rella) grossum, ioannides (1986, p. 16) stated: “archae opyle of dubious position, possibly precingular, type p (or 2p). paracingulum often indicated by aligned processes, which may be joined proximally….” it is confirmed here that the archaeopyle is precingular, resulting from the loss of the 3´´ plate. the operculum is free. chlamydophorella has an apical archaeopyle, so this species is transferred to spongodinium. m. pearce (personal communication 2015) has drawn our attention to the similarity between spongo dinium grossum and isabelidinium? extremum. the latter species has a precingular archaeopyle and an outer wall or periphragm that forms an apical horn and a discontinuous pericyst. thus, this species should be retained as spongodinium? extremum, a combination originally pro posed by lentin & williams (1976). spongodinium grossum and spongodinium? extremum may be synonymous; the only obvious difference seems to be in the nature of the pericyst, which is more continuous and less perforate in spongodinium grossum. spongodinium obscurum (manum & cookson 1964) comb. nov. (plate 13, fig. 19) 1964 scriniodinium obscurum manum & cookson, p. 21, 22, plate 4, figs 5, 6. 1978 scriniodinium? obscurum (manum & cookson) – stover & evitt, p. 188. 2003 endoscrinium obscurum (manum & cookson) – riding & fensome, p. 23. age. lo: santonian. remarks. as demonstrated by ioannides (1986, plates 5–8), intergradation exists between spongodinium (as endoscrinium) obscurum, spongodinium grossum (as chla my dophorella grossa) and a taxon that he called spon go d inium sp. in their description of endoscrinium obscurum, manum & cookson (1964, p. 21) stated: “there are no distinct connections running between the capsule and the theca, but fine, irregular lines, the true nature of which is obscured by the many folds, may possibly represent supporting fibrils. the wall of the capsule is up to 1 μm thick and sometimes indistinctly dotted.” the comments of manum & cookson (1964) and ioannides (1986), together with observations in this study, indicate that the species should be assigned to spongodinium, and that the wall is best considered a complex autophragm, rather than being holocavate. genus stichodinium williams et al. 2015 type. he chengquan & wang kede 1990, plate 2, fig. 3, as wilsonidium subtile. 2015 stichodinium williams et al., p. 314. remarks. williams et al. (2015) erected the genus sticho dinium for wetzelielloidean cysts with a latiepeliform archaeopyle and sutural or penitabular ornamentation that may be features of low relief or processes that are distally free. stichodinium lineidentatum (deflandre & cookson 1955) williams et al. 2015 (plate 13, figs 17, 18) 1955 wetzeliella lineidentata deflandre & cookson, p. 253, 254, plate 5, fig. 5; text-figs 17, 18. 1976 wilsonidium lineidentatum (deflandre & cookson) – lentin & williams, p. 139. 2015 stichodinium lineidentatum (deflandre & cookson) – williams et al., p. 314. age. lo: lutetian. remarks. the holotype of stichodinium lineidentatum (deflandre & cookson 1955, plate 5, fig. 5) clearly shows a latiepeliform archaeopyle. genus subtilisphaera jain & millepied 1973 type. jain & millepied 1973, plate 3, fig. 31, as sub tilisphaera senegalensis. 1973 subtilisphaera jain & millepied, p. 26, 27. remarks. our concept of the genus subtilisphaera adheres to the synopsis presented in fensome et al. (2009), who acknowledged uncertainty in interpreting the archaeopyle. lentin & williams (1976, p. 118) considered that, where bulletin36.qxp_bulletin 36 19/12/16 13.40 side 69 70 observable, subtilisphaera has a combination archaeopyle, with the formula aip(3´+1–3a+3–5´´), the operculum remaining attached along the posterior margin. stover & evitt (1978, p. 238) considered the archaeopyle of subtili sphaera to be “presumably intercalary”. bujak & davies (1983, p. 62) observed specimens from the early cre taceous of offshore eastern canada, which they attributed to subtilisphaera, with complete or incomplete archaeopyle sutures between the following plates: 2´/1a, 3´/1a, 3´/2a, 3´/3a. 4´/3a and with complete or incomplete archaeopyle sutures between individual intercalary plates. these authors termed this a transverse archaeopyle. subtilisphaera perlucida (alberti 1959) jain & millepied 1973 (plate 14, fig. 5) 1959 deflandrea perlucida alberti, p. 102, plate 9, figs 16, 17. 1959 deflandrea pirnaensis alberti, p. 100, plate 8, figs 1, 5. 1960 scriniodinium cooksoniae anderson, p. 30, plate 9, figs 1–3. 1973 subtilisphaera perlucida (alberti) – jain & millepied, p. 27. 1973 subtilisphaera pirnaensis (alberti) – jain & millepied, p. 27. 1990 subtilisphaera? pirnaensis (alberti) – harker & sarjeant in harker et al., p. 133. age. lo: albian. remarks. fensome et al. (2009, p. 61, 62) considered subtilisphaera pirnaensis to be a taxonomic junior syno nym of subtilisphaera perlucida. genus surculosphaeridium davey et al. 1966 emend. nov. type. sarjeant 1960, plate 6, fig. 2, as hystricho sphaeridium cribrotubiferum. 1966 surculosphaeridium davey et al., p. 160, 161. emended diagnosis. chorate gonyaulacacean cysts with spheroidal central bodies. acavate. processes are solid and branched and/or distally furcate. there may be one to four processes per plate, sometimes forming com plexes in cases where there is more than one process per plate. archaeopyle apical, with the formula a(1–4´); operculum free. remarks. surculosphaeridium is emended to allow inclusion of forms with more than one process per plate. however, there must be processes or process complexes on all the plates, including the precin gulars. surculosphaeridium convocatum sp. nov. (plate 14, figs 6–8) holotype. plate 14, fig. 7, from a cuttings sample at 3510–3520 m in south labrador n-19, gsc type collection no. 138128, sample p39834, slide 01, coordinates 7.8 × 99.1, england finder g29/3‒h28/2. central body length (without apical operculum) 30 μm, central body width 38 μm, processes up to about 12 μm. the sample from which the holotype derives is dated as barremian‒aptian, indicating that the specimen represents caving. etymology. the epithet is from the latin convocatum, to call together or assemble, in reference to the grouping of several processes per plate common in this species. diagnosis. a species of surculosphaeridium in which at least the larger plates have more than one and up to four processes per plate. processes may branch along their length and are always furcate distally. size. central body length (without apical operculum) 29‒34 μm, central body width 30‒38 μm, processes up to about 15 μm; three specimens measured. age. lo: early campanian. not plotted. remarks. some plates, especially smaller ones such as cingulars, sulcals and apicals, may have only one process per plate, but others have two to four. processes are characteristically furcate distally. surculosphaeri dium lon gi furcatum has one process per plate. genus talladinium williams, damassa, fensome & guerstein in fensome et al. 2009 type. mao shaozi & norris 1988, plate 13, fig. 6, as charlesdowniea wulagense. 2009 talladinium williams, damassa, fensome & guerstein in fensome et al., p. 61, 62. remarks. williams, damassa, fensome & guerstein in fensome et al. (2009) erected the genus talladinium for wetzelielloidean cysts with a soleiform archaeopyle bulletin36.qxp_bulletin 36 19/12/16 13.40 side 70 71 and processes that are distally united by ectophragmal membranes that delineate the tabulation. talladinium? clathratum (eisenack 1938) williams, damassa, fensome & guerstein in fensome et al. 2009 (plate 14, figs 9, 10) 1938 wetzeliella clathrata eisenack, p. 187; text-fig. 5. 1976 kisselevia? clathrata (eisenack) – lentin & williams, p. 136. 1989 charlesdowniea clathrata (eisenack) – lentin & vozzhennikova, p. 227. 2009 talladinium? clathratum (eisenack) – williams, damassa, fensome & guerstein in fensome et al., p. 62. age. lo: latest bartonian. not plotted. remarks. as noted in fensome et al. (2009, p. 62), the nature of the periand endoarchaeopyles in talladi nium? clathratum can not be ascertained from the original description of eisenack (1938) or the expanded description of eisenack (1954). from the late eocene to early oligocene age of the type material, however, it seems reasonable to deduce that this species has a soleiform archaeopyle. this explains why the species is questionably included in talladinium. talladinium pellis sp. nov. (plate 14, figs 11, 12, 16) holotype. plate 14, figs 11, 12, from a cuttings sample at 1840 m in gjoa o-37, gsc type collection no. 138082, sample yd16082, slide 03, coordinates 050 × 0907, england finder e20/3. size: pericyst length 94 μm, width 76 μm; endocyst length 61 μm, width 57 μm. the age determined for the sample from which the holotype was recovered is late ypresian. etymology. the epithet is from the latin pellis, meaning skin, in reference to the ectophragm, which surrounds most of the pericyst. it is a noun in apposition. diagnosis. a species of talladinium in which processes of individual plates are distally united by a membrane that mimics the outline of the underlying plate. the ectophragm may be irregularly or regularly perforate. size. pericyst length 89‒94 μm, width 76‒81 μm; endocyst length 61 μm (both specimens), width 57‒64 μm; two specimens measured. age. lo: priabonian. remarks. the ectophragmal membranes in talladi nium pellis form linear complexes that mirror the outline of the reflected plates, rather than forming a shield-like covering as in talladinium? clathratum. this difference is not always easy to determine. charlesdowniea coleothrypta has identical ornamentation but differs in having an equiepeliform archaeopyle. genus tanyosphaeridium davey & williams 1966a type. davey & williams 1966a, plate 6, fig. 7, text-fig. 20, as tanyosphaeridium variecalamum. 1966a tanyosphaeridium davey & williams, p. 98. remarks. tanyosphaeridium is characterised by an elongate ellipsoidal central body, slender open processes and an apical archaeopyle. the processes number about 30 or more and cannot be readily related to tabulation. tanyosphaeridium xanthiopyxides (wetzel 1933a ex deflandre 1937) stover & evitt 1978 (plate 15, fig. 7) 1933a hystrichosphaera xanthiopyxides wetzel, p. 44, 45, plate 4, fig. 25 (name not validly published). 1937 hystrichosphaeridium xanthiopyxides wetzel ex deflandre, p. 77. 1965 baltisphaeridium xanthiopyxides (wetzel ex deflandre) – downie & sarjeant, p. 98. 1968 hystrichosphaeridium? xanthiopyxides (wetzel ex deflandre) – morgenroth, p. 556. 1969 prolixosphaeridium? xanthiopyxides (wetzel ex deflandre) – davey et al., p. 17. 1978 tanyosphaeridium xanthiopyxides (wetzel ex deflandre) – stover & evitt, p. 85. age. lo: danian. remarks. fensome et al. (2009, p. 62) is followed here in using the earliest proposed name, tanyosphaeridium xanthiopyxides, for all the late cretaceous – paleocene forms of tanyosphaeridium. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 71 72 genus taurodinium gen. nov. type.  plate 14, fig. 13, as taurodinium granulatum. etymology. the name derives from the latin taurus, mean ing bull, in reference to the two prominent apical horns. description. dorso-ventrally compressed ceratiacean cysts with an autocyst having six horns: two apical, two lateral and two antapical, all with closed, acuminate to slightly rounded terminations. autophragm smooth to finely granulate or perforate and very thin. parata bula tion not expressed. cingulum may be indicated above the two lateral horns but there is no indication of a sulcus. archaeopyle apical with a straight to weakly angular margin; operculum usually attached. remarks. nyktericysta and vesperopsis have the same thin and fragile appearance as taurodinium, but have only a single apical horn. satyrodinium lentin & ma num 1986, has two to three apical and one or more antapical horns, but is a peridiniacean cyst with an intercalary archaeopyle, lacks lateral horns, and is cavate. the acritarch genus limbicysta marshall 1989, from upper cretaceous non marine to nearshore environments, has an elongate split protrusion at one end that shows some resemblance to the apical horns of taurodinium but limbicysta lacks a regular opening and other protrusions. taurodinium granulatum sp. nov. (plate 14, figs 13–15, 17, 18) 1992 gen. et sp. indet. piasecki et al., fig. 6l, m. 1995 nyktericysta sp. gregory & hart, plate 7, figs 6, 7. 2011 gen. et sp. indet. piasecki et al. 1992 – nøhrhansen et al., figs 3, 4. 2012 gen. et sp. indet. piasecki et al. 1992 – nøhrhansen, plate x, figs 12–19. holotype. plate 14, fig. 13, from a cuttings sample at 2340 m in ikermiut-1, offshore west greenland, mghu no. 31333, sample ikermiut 2289, slide s-261-3, england finder t19-1. size: overall length 105 μm, length of body 54 μm, body width 35 μm, length of apical horns up to 30 μm, length of lateral horns 13 μm, length of antapical horns 22 μm, wall less than 1 μm thick. the age determined for the sample from which the holotype was recovered is late thanetian. etymology. the epithet is from the greek granulatum meaning granulate, in reference to the surface ornamentation. description. a species of taurodinium with two apical horns of approximately equal length, two short lateral horns and two antapical horns of unequal length with the left being the longer. the antapical horns narrow at about one-third along their length. autophragm smooth to finely granulate, thin, hyaline and often wrinkled. the apical horns appear to be oriented in a plane perpendicular to that of the other horns, giving the cyst a twisted look. size.  overall length 75–105 μm, length of body 40–70 μm, body width 28–60 μm, length of apical horns 15– 25 μm, length of lateral horns 7–17 μm, length of antapical horns 10–32 μm; seven specimens measured. age. lo: bartonian; peak early ypresian. remarks. taurodinium granulatum was originally record ed as ‘gen. et sp. indet.’ by piasecki et al. (1992) from a miospore-dominated palynological assemblage devoid of other dinocysts. the assemblage was from a silty shale clast in subaqueous volcanic breccias from the lower rinks dal member, maligât formation, west green land; the basalts were radiometrically dated as 61.2 ± 0.4 ma, and thus of selandian age (larsen et al. 2015). gregory & hart (1995) recorded a specimen of ny k tericysta sp. that may be assignable to taurodinium gra nulatum from sediments dated as thanetian. nøhrhansen et al. (2011) recorded gen. et sp. indet. of piasecki et al. (1992) from thanetian – lower? ypresian non-marine or marginal marine deposits of north-eastern greenland. recently, this species has been found to be common in the thanetian‒ypresian non-marine or marginal marine kulhøje member in the kangerlussuaq basin, south-eastern greenland (nøhr-hansen 2012). in the present study, taurodinium granulatum was encountered in wells on the west greenland continental margin in samples together with the late paleocene marker axiodinium augustum, as well as in samples of ypresian age from the saglek basin. however, the species has been mostly recorded from spore-dominated or marginal marine, algae-dominated assemblages. we thus follow nøhr-hansen (2012) in regarding it as a freshwater to brackish-water indicator. genus tenua eisenack 1958 type. eisenack 1958, plate 23, fig. 1, as tenua hystrix. 1958 tenua eisenack, p. 410. 1981 cerbia below, p. 8. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 72 73 remarks. the genus is distinguished from cyclone p he lium (incorporating circulodinium) by having some processes showing alignment in penitabular rows. aptea differs in having three horns – an apical, an antapical, and a third, probably postcingular – of variable development; the probable postcingular horn is always re duced. cerbia is considered here to be a taxonomic junior synonym of tenua. tenua hystrix eisenack 1958 (plate 14, figs 19, 20) 1958 tenua hystrix eisenack, p. 410, plate 23, figs 1–4; text-fig. 10. 1972 tenua hystricella eisenack & kjellström, p. 1039. 1978 cyclonephelium hystrix (eisenack) – davey, p. 894. age. lo: late aptian. remarks. sarjeant (1985b, p. 94, 95) retained this species in tenua. the distribution of processes in this species is variable. the material in this study includes forms in which processes are largely absent on midventral and mid-dorsal surfaces (for example that shown in plate 14, fig. 19). genus thalassiphora eisenack & gocht 1960 type. eisenack 1954, plate 12, fig. 17, as pterosper m opsis pelagica. 1960 thalassiphora eisenack & gocht, p. 513. 1966a erikania morgenroth, p. 27. 1980 subathua khanna & singh, p. 307, 308. remarks. fensome et al. (2009, p. 62) provided a com prehensive synopsis for thalassiphora based on the emendations of the genus by williams & downie (1966a, p. 234), gocht (1968, p. 153) and benedek & gocht (1981, p. 59). thalassiphora delicata williams & downie 1966a (plate 15, fig. 1) 1966a thalassiphora delicata williams & downie, p. 235, plate 26, fig. 8. 1973 disphaeria delicata norvick, p. 43. age. lo: middle bartonian. not plotted. remarks. lentin & williams (1977a, p. 54) retained this species in thalassiphora. thalassiphora fenestrata liengjarern et al. 1980 (plate 15, figs 2–4) 1980 thalassiphora fenestrata liengjarern et al., p. 489, plate 54, fig. 1. age. lo: earliest rupelian. remarks. thalassiphora fenestrata has large fenestrations, which are restricted to lateral and ventral areas of the pericyst. liengjarern et al. (1980, p. 489) considered its stratigraphic range to be late eocene to early? oligocene. thalassiphora pelagica (eisenack 1954) eisenack & gocht 1960 (plate 15, figs 5, 6) 1954 pterospermopsis pelagica eisenack, p. 71, plate 12, figs 17, 18. 1960 thalassiphora pelagica (eisenack) – eisenack & gocht, p. 513, 514. 1966 thalassiphora sueroi pöthe de baldis, p. 224, 225, plate 2, fig. d. 1973 disphaeria pelagica (eisenack) – norvick, p. 46. 1981 disphaeria sueroi (pöthe de baldis) – yun hyesu, p. 70. age. lo: chattian. remarks. lentin & williams (1977a, p. 54) retained this species in thalassiphora. genus trichodinium eisenack & cookson 1960 type. eisenack & cookson 1960, plate 2, fig. 4, as trichodinium pellitum. 1960 trichodinium eisenack & cookson, p. 5. remarks. trichodinium is a spheroidal to ovoidal proximate gonyaulacacean cyst that apically can have a horn or several spines. the autophragm may be tabulate, as determined from the alignment of some of the nume rous short spines or bifid processes. its surface has been described as fibro-pitted (see fensome et al. 2009, p. 64). the archaeopyle is precingular, with the formula p3´´; the operculum is free. fensome et al. (2009, p. 64) bulletin36.qxp_bulletin 36 19/12/16 13.40 side 73 74 discussed the morphological similarity between tricho dinium and xenicodinium klement 1960. they recommended that xenicodinium should be restricted to its type material and all forms with the appropriate mor phology should be included in trichodinium. trichodinium castanea deflandre 1935 ex clarke & verdier 1967 (plate 15, fig. 8) 1935 palaeoperidinium castanea deflandre, p. 229, plate 6, fig. 8; name not validly published. 1967 trichodinium castanea (deflandre) – clarke & verdier, p. 19, 20. age. lo: campanian. remarks. some tabulation can usually be discerned on trichodinium castanea, which is also distinguished by its dense covering of short, acuminate to bifid spines. genus trithyrodinium drugg 1967 type. drugg 1967, plate 3, fig. 2, as trithyrodinium evittii. 1967 trithyrodinium drugg, p. 20. synopsis. peridiniacean (deflandreoid) cysts that are proxi mate and rounded to peridinioid in shape, with an antapex that is rounded, symmetrical or, more usually weakly to strongly asymmetrical, the left side being larger; endophragm rounded pentagonal to subcircular in dorso-ventral outline. cavate; endocyst sometimes strongly developed with a fragile or thin-walled peri phragm that is easily lost. archaeopyle intercalary, with formula i(1–3a); operculum free, compound. plate 2a is always isoto steno deltaform. remarks. the synopsis is largely a repeat of that provided by fensome et al. (2009, p. 64), though with the concept broadened to include rounded forms and to recognise that the pericyst may be durable. the 2a plate, which is revealed through archaeopyle development, can show some variation in width but is always deltaform. the genus pierceites also has a 3i archaeopyle but the endocyst is absent or weakly developed. trithyrodinium? conservatum sp. nov. (plate 15, figs 9–14) 2003 deflandrea sp.1 nøhr-hansen, plate 3, figs 4–6. holotype. plate 15, fig. 12 and nøhr-hansen (2003, plate 3, fig. 6), from a sidewall-core sample at 1155 m in ikermiut 1, mghu no. 26502, sample 04e006504, slide 2, co-ordinates 36.0 × 97.1, england finder u36/1. pericyst length 60 μm, width 63 μm, endocyst length 44 μm, width 52 μm. the age determined for the sample from which the holotype was recovered is lutetian. etymology. the epithet is from the latin conservatus, meaning retain or conserve, in reference to the constant presence of the pericyst. description. a species of trithyrodinium with a com mon ly rounded but sometimes ovoidal pericyst that is always present. the endocyst generally mimics the shape of the pericyst. as a rule the cyst is circumcavate, but the endocyst may occasionally be in partial contact with the pericyst. the periphragm and endophragm are both thin, at the most slightly over 1 μm thick. the periphragm varies from laevigate (the usual condition), to faintly granulate or verrucate. the i to 3i archaeopyle is formed from the loss of one to three intercalary plates individually, any of which can remain attached posteriorly. size. pericyst length 48–60 μm, width 49–65 μm, endo cyst length 44–51 μm, width 43–56 μm; seven specimens measured. age. lo: lutetian. not plotted. remarks. trithyrodinium? conservatum is unusual in that the pericyst is always preserved. when the pericyst has a sub dued granulate or verrucate ornamentation, this tends to be restricted to the mid-dorsal and mid-ventral regions. folds are consistently present on the pericyst but appear to be random. the exact nature of the archaeopyle is unclear. in some specimens its polygonal shape appears to indicate that multiple plates are missing, but in others it appears to reflect loss of a single intercalary plate. because of the uncertainty regarding the archaeopyle, the species is only assigned questionably to trithyrodinium. trithyrodinium evittii drugg 1967 (plate 15, figs 17–19) 1967 trithyrodinium evittii drugg, p. 20, plate 3, figs 2, 3; plate 9, fig. 2. 1969b trithyrodinium fragile davey, p. 11, plate 3, figs 6, 9. age. lo: danian. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 74 75 trithyrodinium quinqueangulare marheinecke 1992 (plate 15, fig. 16) 1992 trithyrodinium quinqueangulare marheinecke, p. 95, plate 19, figs 9–11. age. lo: maastrichtian. remarks. trithyrodinium quinqueangulare has a pericyst and endocyst, which are sometimes closely appressed. both pericyst and endocyst are pentagonal and have two more or less equal antapical protuberances. the cingulum is clearly delineated. plate 4´´ is reduced in an anterior–posterior direction. some of the labrador margin specimens have a verrucate endophragm. trithyrodinium suspectum (manum & cookson 1964) davey 1969b (plate 15, fig. 15) 1964 hexagonifera suspectum manum & cookson, p. 9, 10, plate 1, figs 9–13. 1969b trithyrodinium suspectum (manum & cookson) – davey, p. 12. age. lo: campanian. remarks. trithyrodinium suspectum has a thick granular endophragm, which appears to be tectate. the archaeo pyle may form from the detachment of the 1a and 3a plates, whereas the 2a plate can remain attached posteriorly, along its boundary with the 4´´ plate. one spe cimen of trithyrodinium suspectum illustrated in ma num & cook son (1964, plate 1, fig. 11) has a steno deltaform 2a plate rather than the more commonly observed isodeltaform 2a. genus tuberculodinium wall 1967 type. rossignol 1962, plate 2, fig. 1, as pterospermopsis? vancampoae. 1967 tuberculodinium wall, p. 114. synopsis. goniodomacean (gambierdiscoid) cysts proximate, preicyst and endocyst both discoidal to subsphe roidal. holocavate. numerous intratabular pillarto barrel-shaped processes support an ectophragm. archae o pyle antapical, involving usually 2–3 paraplates, operculum free, compound. remarks. tuberculodinium represents the cysts of pyro phacus, with its multiplate tabulation; the ectophragmal supports (processes), although numerous, can be readily interpreted as intratabulate. tuberculodinium vancampoae (rossignol 1962) wall 1967 (plate 15, fig. 20) 1962 pterospermopsis? vancampoae rossignol, p. 134, plate 2, fig. 1. 1967 tuberculodinium vancampoae (rossignol) – wall, p. 114, 115. 1971 pyrophacus vancampoae (rossignol) – wall & dale, p. 234. age. lo: tortonian? remarks. head (1996, p. 1232) retained this species in tuberculodinium. genus vesperopsis bint 1986 type. bint 1986, plate 5, figs 9, 12–13; text-fig. 5, as vesperopsis mayii. 1986 vesperopsis bint, p. 156. remarks. vesperopsis is a proximate ceratiacean cyst with an autophragm that has at least three horns – one apical and two antapical. commonly, it may also have equatorial horns with or without preand postcingular branches. the archaeopyle is apical, with the formula a(1–4´); the operculum is usually attached. vesperopsis was emended by qiao xiuyun et al. (1992, p. 32, 33, 36, 37), wan chuanbiao & qiao xiuyun (1994, p. 503) and mao shaozhi et al. (1999, p. 149, 150). however, we adhere to the synopsis of fensome et al. (2009, p. 65). vesperopsis longicornis (batten & lister 1988) harding 1990 (plate 16, figs 1, 2) 1988 australisphaera longicornis batten & lister, p. 340, 341, fig. 1b–e, g. 1990 vesperopsis longicornis (batten & lister) – harding, p. 21. age. lo: albian. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 75 76 remarks. harding (1990, p. 21) emended the diagnosis of vesperopsis longicornis, based on it having an auto phragm and an attached operculum. according to har ding, an unusual aspect of vesperopsis longicornis is that the operculum appears to be attached dorsally; operculum attachment in ceratiacian cysts, where present, is almost always ventral. genus wallodinium loeblich jr. & loeblich iii 1968 type. cookson & eisenack 1960b, plate 39, fig. 4, as diplotesta glaessneri. 1960b diplotesta cookson & eisenack, p. 256; name illegitimate. 1968 wallodinium loeblich jr. & loeblich iii, p. 212. synopsis. proximate gonyaulacalean cyst, elongate crescent-shaped to subcylindrical. bicavate, circumcavate or epicavate. pericyst much longer than endocyst; both bodies generally rounded apically and antapically, al though the endocyst may have short horns. tabulation indicated by apical archaeopyle, with inferred formula of a(1–4´); periand endoarchaeopyle free or attached. sometimes an equatorial constriction or ornamentation alignment marks a cingulum. wall smooth or with orna mentation of low relief. remarks. wallodinium was considered to be an acritarch by duxbury (1983, p. 68) and fensome et al. (1990, p. 535). however, riding (1994, p. 17, 18) emended the diagnosis to include reference to the indications of gonyaulacalean tabulation and the presence of an apical archaeopyle; the above synopsis is based on riding’s emendation. wallodinium luna (cookson & eisenack 1960a) lentin & williams 1973 (plate 16, fig. 5) 1960a diplotesta luna cookson & eisenack, p. 10, 11, plate 3, fig. 21. 1973 wallodinium luna (cookson & eisenack) – lentin & williams, p. 140. age. lo: campanian. genus wetzeliella eisenack 1938 emend. williams, damassa, fensome & guerstein in fensome et al. 2009 type. eisenack 1938, fig. 4, as wetzeliella articulata. 1938 wetzeliella eisenack, p. 187. 1979 gochtodinium bujak, p. 310–312. 2009 wetzeliella eisenack – emend. williams, damassa, fensome & guerstein in fensome et al., p. 65. synopsis. wetzelielloidean cysts with a soleiform archae o py le and processes that are predominantly non tabular, al though they can show some alignment, and are dis tally free. wetzeliella articulata wetzel in eisenack 1938 emend. williams, damassa, fensome & guerstein in fensome et al. 2009 (plate 16, figs 3, 4) 1938 wetzeliella articulata wetzel in eisenack, p. 187, text fig.4. 1938 palaeoperidinium articulatum wetzel in eisenack, p. 187 (name not validly published). 1948 hystrichosphaeridium articulatum (wetzel in eisenack) – pastiels, p. 42. 1960 wetzeliella echinulata vozzhennikova, plate 3, fig. 3 (name not validly published). 1967 rhombodinium coronatum vozzhennikova, p. 170, 171, plate 89, figs 1–3, 5; plate 90, figs 1–5. 1967 wetzeliella echinulata vozzhennikova, p. 164, 165. 1975 wetzeliella horrida jan du chêne & châteauneuf, p. 28, 30, plate 1, figs 1–7; plate 3, figs 1–6. 1976 wetzeliella coronata (vozzhennikova) – lentin & williams, p. 131. age. lo: rupelian. not plotted. remarks. according to the literature, wetzeliella articulata has a stratigraphic range of eocene–oligocene. but when the species is restricted to wetzeleilloideans with a soleiform archaeopyle, intratabular processes and five horns, its stratigraphic range is more restricted, being bartonian to rupelian. genus xenascus cookson & eisenack 1969 type. cookson & eisenack 1969, fig. 1i, j, as xenascus australensis. 1969 xenascus cookson & eisenack, p. 7. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 76 77 remarks. fensome et al. (2009, p. 67) provided a com prehensive synopsis of xenascus in which they referred to lateral rather than postcingular, horn(s) because they seem to emanate from the cingulum as well as postcingular areas. fensome et al. (2009) also discussed the emendations of yun hyesu (1981, p. 60) and stover & helby (1987, p. 128). xenascus ceratioides (deflandre 1937) lentin &williams 1973 (plate 16, figs 9, 10) 1937 hystrichosphaera ceratioides deflandre, p. 66, 67, plate 12 (also labelled as plate 9), figs 7, 8. 1967 pseudoceratium ceratioides (deflandre) – clarke & verdier, p. 60. 1970 spiniferites ceratioides (deflandre) – sarjeant, p. 76. 1971 phoberocysta ceratioides (deflandre) – davey & verdier, p. 26. age. lo: campanian. remarks. xenascus ceratioides has three horns: one api cal, one lateral and one antapical. the horns are not perforated. genus fromea (cookson & eisenack 1958) yun hyesu 1981 type. cookson & eisenack 1958, plate 5, fig. 10, as fromea amphora. 1958 fromea cookson & eisenack, p. 55. 1973 xenascus ceratioides (deflandre) – lentin & williams, p. 144. remarks.  although many workers assign fromea to the dinoflagellates, some (e.g. duxbury 1980, p. 134; fen some et al. 1990, p. 227; fensome & williams 2004, p. 742) consider it to lack unequivocal morphological indications of dinoflagellate affinity and thus prefer to consider it an acritarch. xenascus wetzelii slimani 1996 ex slimani 2001a (plate 16, figs 6–8) 1985 odontochitina wetzelii wilson in foucher in robaszynski et al., p. 33, plate 10, figs 9–12; name not validly published. 1996 xenascus wetzelii slimani, p. 380, 381, plate 3, figs f, g; plate 4, figs a, b; text-fig. 7a, b; name not validly published. 2001a xenascus wetzelii slimani 1996 ex slimani, p. 9, plate 2, figs 3, 4. age. lo: campanian. remarks. xenascus wetzelii is cornucavate to circumcavate. it has a long postcingular horn and a long antapical horn, both of which are perforate distally and can be acuminate or bifurcate at their extremities. tabulation is clearly shown by parasutural crests and gonal processes, the latter being acuminate, bifurcate or trifurcate. slimani (2001a, p. 9; 2001b, p. 194) considered odon tochitina wetzelii to be a taxonomic junior synonym of xenascus wetzelii. fromea nicosia jansonius 1989 (plate 16, fig. 19) 1989 fromea nicosia jansonius, p. 67, plate 1, figs 2–7; text-fig. 1. age. lo: early campanian. fromea quadrangularis sp. nov. (plate 16, figs 11, 12, 15, 16) holotype. plate 16, fig. 16, from a cuttings sample at 2375 m in skolp e-07, mguh no. 31352, sample yd15665, slide 04, co-ordinates 30.9 × 111.9, eng land finder d31-1. overall length 88 μm; maximum width 35 μm; minimum width 27 μm; wall thickness 1 μm. systematics – acritarchs and other algae bulletin36.qxp_bulletin 36 19/12/16 13.40 side 77 78 the age determined for the sample from which the holotype was recovered is early campanian, late cretaceous. etymology.  the epithet is from the latin words quadra, meaning four and angularis, having angles, in reference to the quadrangular shape of specimens of this species. diagnosis.  a slender elongate rectangular species of fro mea with a thin smooth wall and generally two longitudinal folds. where present, the folds start at the antapex and continue along the body to the circular ‘apical’ opening. no accessory opening sutures or equatorial ‘girdle’ have been observed. size.  overall length 77–101 μm; maximum width 30– 45 μm; minimum width 27–37 μm; seven specimens measured. age.  lo: early campanian. remarks.  some specimens of fromea quadrangularis are twisted, resulting in an elongate jar-like shape with longitudinal folds; other specimens have been slightly compressed, resulting in a weak equatorial extension and a splitting of the elongate folds. genus microsphaeridium benedek 1972 type. benedek 1972, plate 12, fig. 3a, b, as micro sphaeridium ancistroides. 1972 microsphaeridium benedek, p. 46, 47. remarks. benedek & sarjeant (1981, p. 346, 347) interpreted microsphaeridium ancistroides as “the de tached opercula of skolochorate dinoflagellate cysts.” pending further investigation, we prefer to consider microsphaeridium to be an acritarch. microsphaeridium ancistroides benedek 1972 (plate 16, figs 13, 14, 17, 18) 1972 microsphaeridium ancistroides benedek, p. 47, plate 12, fig. 3a, b; text-fig. 21. age. lo: middle miocene or younger; not well constrained. not plotted. genus palambages wetzel 1961 type. wetzel 1961, plate 1, fig. 11, as palambages mo ru losa. 1961 palambages wetzel, p. 338. remarks. wetzel (1961) noted that the microfossils as signed to palambages are identical to forms recorded as “morulosae” in wetzel (1933a, p. 23, 24, plate 4, figs 1–5). wetzel (1933a) had compared these fossils with certain colonial algae, but noted their likeness to the egg-balls of planktonic crustaceans. manum & cook son (1964, p. 23) considered them to represent colonies of green algae. palambages spp. (plate 16, fig. 20) age. range not determined. not plotted. genus paralecaniella cookson & eisenack 1970a type. deflandre & cookson 1955, plate 9, fig. 6, as epicephalopyxis indentata. 1970a paralecaniella cookson & eisenack, p. 323. remarks. understanding of this genus, which is abundant in what seem to be shallow-water environments, is tenuous. elsik (1977, p. 96), who emended the diagnosis of the genus and the type, considered paralecaniella to be a dinocyst. we agree with fensome et al. (2009, p. 67) that it is probably an algal cyst of non-dinoflagellate affinity. paralecaniella indentata (deflandre & cookson 1955) cookson & eisenack 1970a (plate 17, figs 1, 2) 1955 epicepholopyxis indentata deflandre & cookson, p. 292, plate 9, figs 5–7; text-fig. 56. 1970a paralecaniella indentata (deflandre & cookson) – cookson & eisenack, p. 323. 1973 scriniodinium? nilsii kjellström, p. 42, fig. 35. age. peak occurrence within the earliest ypresian. remarks. the peak occurrence of paralecaniella indentata is close to the paleocene–eocene boundary. genus pediastrum meyen 1829 type. meyen 1829, plate 43, figs 6–20, as pediastrum duplex. 1829 pediastrum meyen, p. 772. remarks. pediastrum is a nonmotile coenobial green alga that is found in freshwater environments. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 78 79 pediastrum spp. (plate 17, figs 3, 4) age. range not determined. not plotted. the range of the genus extends from early cretaceous to recent (batten 1996). remarks. the presence of pediastrum spp. in some of the samples from marine deposits in the offshore wells is indicative of offshore transport of freshwater elements. genus tetraporina naumova 1939 type. tetraporina antiqua naumova 1950, designated by potonié (1960, p. 130). 1939 tetraporina naumova, p. 357. 1956 tetrapidites klaus in meyer, p. 107. 1960 tetraporopollenites frantz, p. 559. 1963 balmeella pant & mehra, p. 116. 1980 tetraporina naumova – emend. lindgren, p. 346. synopsis. acid resistant, unicellular, tetrahedral or par al lel epipedal microfossils with or without obvious pore or other dehiscence mechanism. wall single or double layered. remarks. described originally as representing pollen, hemer & nygreen (1967) considered tetraporina to be an algal genus, and it has been considered generally to represent the acritarchs or algae since then. the above synopsis largely follows the emended diagnosis of lindgren (1980). recognition of tetaporina is based mainly on its distinctive tetrahedral or parallelepipedal shape (a parallelepiped is a geometric figure with six faces, all parallelograms and all opposite faces being similar and parallel). although most commonly related to modern zygospores of the zygnemataceae, lindgren (1980) considered teraporina to be polyphyletic and difficult to match with particular modern algae because of its simple shape, rigours of preservation and ontogenetic factors. most algae to which tetraporina can be related are freshwater. the stratigraphic range of the genus extends from carboniferous to quaternary. there has been considerable debate about the nomenclatural status of tetraporina. jansonius & hills (1981, card 3917) considered that the name was not validly published in naumova (1939) because it was proposed in anticipation of future acceptance of the name (see mcneill et al. 2012, article 36.2). however, the case for this has not proved compelling for other authors (lindgren 1980; farr & zijlstra 1996). although both sets of arguments have merit, for pragmatic reasons, here we accept tetraporina as valid in naumova (1939). tetraporina sp. a (plate 17, figs 5–7) description. a form of tetraporina with a tetrahedral outline and indentations at the apices, although these are not clearly perforated by pores. the apices have no, or only short, extensions. the wall surface is smooth or has subdued ornament such as small verrucae. this form may correspond with one or more of the species described by lindgren (1980), but morphological overlap, preservation and flexibility of the wall, and the quality of lindgren’s illustrations make comparison difficult. age. lo: earliest rupelian. tetraporina? sp. b (plate 17, figs 8–12) description. a form with broad truncated extensions at each of the four apices. the extensions may be relatively short and converge on a central rhombic area, or they may just converge centrally without the development of a central rhombic area, in which case the entire specimen is essentially a cross-like structure. tetraporina does not typically have extensions, so this form is only tentatively associated with that genus. age. not plotted. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 79 80 miospores genus afropollis doyle et al. 1982 type. brenner 1968, plate 10, fig. 5, as reticulato sporites jardinus. 1982 afropollis doyle et al., p. 44. afropollis sp. (plate 17, figs 18, 19) remarks. this form is assigned to afropollis because of its possession of a network of reticulate to regulate muri separated from the nexine. the grains are apparently inaperturate. no ring furrow was discerned. the net work of muri is considerably coarser than that exhibited by afropollis jardinus, afropollis operculatus and afro pollis zonatus. age. lo: cenomanian. genus appendicisporites weyland & krieger 1953 type. weyland & krieger 1953, plate 11, fig. 54, as appendicisporites tricuspidatus. 1953 appendicisporites weyland & krieger, p. 12. appendicisporites potomacensis brenner 1963 (plate 17, fig. 20) 1963 appendicisporites potomacensis brenner, p. 46, plate 6, figs 4, 5. 1985 plicatella potomacensis (brenner) – davies, p. a49. age. lo: mid-cretaceous, but not well constrained. not plotted. remarks. this species is retained in appendicisporites, following burden & hills (1989, p. 100) and nichols & sweet (1993, p. 548). appendicisporites unicus (markova in ivanova & markova 1961) singh 1964 (plate 18, fig. 1) 1961 anemia unica markova in ivanova & markova, p. 53, plate 20, fig. 3a, b. 1964 appendicisporites unicus (markova in ivanova & markova) – singh, p. 53. 1985 plicatella unica (markova in ivanova & markova) – davies, p. a 53. age. lo: mid-cretaceous, but not well constrained. not plotted. remarks. this species is retained in appendicisporites, following burden & hills (1989, p. 100). genus aquilapollenites rouse 1957 emend. braman 2013 type. radforth & rouse 1954, plate 1, fig. 14, as “n2”. holotype lost; lectotype (possibly the holotype) selected from restored type slide by tschudy & leopold (1971, plate 2, fig. 1); neotype designated by srivastava & rouse (1970, plate 1, figs 4–7); all as aquilapollenites quadrilobus. 1957 aquilapollenites rouse, p. 370. 1970 hemicorpus krutzsch, p. 107. 2013 aquilapollenites rouse – emend. braman, p.14. remarks. braman (2013) is followed here in considering aquilapollenites to comprise triprojectate pollen that are heteropolar and have isolated sculptural ele ments. other triprojectate pollen found in this study are assignable to parviprojectus and translucentipollis. aquilapollenites quadrilobus rouse 1957 (plate 17, figs 14, 15, 17) 1957 aquilapollenites quadrilobus rouse, p. 371, plate 2, figs 8, 9. 1961 aquilapollenites polaris funkhouser, p. 198, plate 1, figs 1, 2. 1961 aquilapollenites pulcher funkhouser, p. 198, plate 1, fig. 7a–c. 1970 mancicorpus polaris (funkhouser) – stanley, p. 30. 1970 hemicorpus polaris (funkhouser) – krutzsch, p. 107. 1970 hemicorpus pulcher (funkhouser) – krutzsch, p. 107. 1970 mancicorpus pulcher (funkhouser) – srivastava, p. 697. systematics – miospores and fungal elements bulletin36.qxp_bulletin 36 19/12/16 13.40 side 80 81 1970 aquilapollenites quadrilobus rouse – emend. srivastava & rouse, p. 1597. remarks. the synonymy above follows that in braman (2013, p. 28). this widely recorded species is characterised by heteropolar grains with prominent spinate sculpture. according to braman (2013) and a. sweet (personal communication 2015), this species has a late cretaceous range, so the specimens recorded from the labrador–baffin seaway are assumed to be reworked. age. lo: middle? eocene (reworked?). genus azolla lamarck in lamarck et al. 1783 type. azolla filiculoides lamarck et al. 1783. 1783 azolla lamarck in lamarck et al., p. 343. azolla spp. (plate 18, figs 2–4) age. lo: ?bartonian. frequent in late ypresian. remarks. azolla is a small moss-like, free-floating freshwater fern that occurs today in warm climates; it is famous for its nitrogen-fixing capability. given the right combination of time and temperature, sediment rich in azolla could represent major sources of oil. azolla had coeval blooms in some arctic and northern temperate areas during the early to earliest middle eocene (e.g. barke et al. 2011, 2012). barke et al. (2012) concluded that the presence of the blooms accords with high-precipitation conditions modelled for the early eocene and implies the presence of extensive wetlands bordering the regional landmasses. they further suggested that azolla blooms in the arctic and norwegian sea basins may indicate widespread fresh ocean surface waters due to unprecedented discharge of water from the land and the semi-enclosed nature of those two basins. the occurrences of azolla in labrador sea – davis strait wells are mostly restricted to a narrow time interval at the top of the ypresian. azolla has not been recorded in the offshore west greenland wells. evidence of extensive blooms of azolla in the lab rador–baffin seaway was not encountered in this study, perhaps indicating that this marine basin was broadly open to the south and had limited or no connection to the arctic basin (see barke et al. 2012, fig. 4 and nøhrhansen et al. 2016). the poor preservation and limited frequency of azolla specimens in our study have precluded specific assigments, but several species are probably represented (see barke et al. 2012). previous studies suggest an lo of azolla species in the earliest lutetian. genus baculatisporites pflug & thomson in thomson & pflug 1953 type. wolff 1934, plate 5, fig. 8, as sporites priarius. 1953 baculatisporites pflug & thomson in thomson & pflug. synopsis. trilete spores, with a subcircular ambitus; the laesurae are of variable length, sometimes reaching to the equator. the ornamentation of the exine is predominantly of bacula, although other types of element may be present. distally, the bacula can be flat to multicrowned. remarks. in our synopsis for baculatisporites, we allow for variation in both the width relative to the height of the baculae, so that the width may exceed the height, and in the nature of their distal terminations. baculatisporites crenulatus sp. nov. (plate 18, figs 6–8) holotype. plate 18, fig. 8, from a cuttings sample at 1790–1800 m in roberval k-92, gsc type collection no. 137889, sample p17684, slide 01, co-ordinates 10.6 × 104.8, england finder l34/2. maximum overall diameter 65 μm. the age determined for the sample from which the holotype was recovered is bartonian. etymology. the epithet is from the latin crenulatus, meaning minutely crenulate, in reference to the toothed distal margin of the bacula. diagnosis. a species of baculatisporites with sculptural elements that have variable distal terminations, ranging from blunt to rounded to crenulate. there is considerable variation in the nature of the distal terminations, and also in the length-to-width ratio of the elements: many are of greater length than width and hence true bacula, but others are of greater width than length. size. diameter 52–65 μm. length of laesurae 15–18 μm. bacula height 3–12 μm, width 3–7 μm. wall thickness 2–3 μm. age. lo: bartonian. not plotted. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 81 82 remarks. the ratio of the width to height of the baculae and their distal variability from blunt to crenulated, distinguishes baculatisporites crenulatus from other spe cies of the genus. genus callialasporites dev 1961 type. balme 1957, plate 8, fig. 91, as zonalapollenites trilobatus. 1961 callialasporites dev, p. 48. 1961 applanopsis döring, p. 112. 1961 triangulopsis döring, p. 113. 1962 pflugipollenites pocock, p. 72. 1964 applanopsipollenites levet-carette, p. 107. 1970 singhiapollis kar & sah, p. 107. synopsis. cavate, proximo-distally compressed mio spores with a roughly circular to triangular or trilobate amb. the two wall layers are mostly appressed proximally and distally, but ambitally and sub-ambitally the wall layers are variously separated to produce a hollow zona that may be continuous or constricted to few (commonly three) to multiple vesicles and/or be radially plicated. the central body may be circular, triangular, or irregularly shaped. a proximal, non-functional triradiate mark may be present. remarks. the above description is condensed from a version provided by fensome (1983), who gave an extensive review of this genus and its synonymy. callialasporites dampieri (balme 1957) dev 1961 (plate 18, fig. 16) 1957 zonalapollenites dampieri balme, p. 32, plate 8, figs 88–90. 1961 callialasporites dampieri (balme) – dev, p. 48. 1961 applanopsis dampieri (balme) – döring, p. 113. 1962 pflugipollenites dampieri (balme) – pocock, p. 72. 1963 tsugaepollenites dampieri (balme) – dettmann, p. 100. remarks. singh (1971, p. 175) retained this species in callialasporites. age. lo: aptian. callialasporites obrutus norris 1969 1969 callialasporites obrutus norris, p. 597, plate 110, figs 6, 7. age. lo: aptian. genus caryapollenites raatz 1938 ex potonié 1960 type. potonié 1931, fig. 2, as pollenites simplex. 1934 caryae?-pollenites potonié & venitz, p. 21; name not validly published. 1938 caryapollenites raatz, p. 19; name not validly published. 1960 caryapollenites potonié, p. 123. remarks. as nichols & ott (1978) noted, the nomenclatural history of fossil pollen similar to the modern genus carya is difficult to unravel. this is also true of species within caryapollenites. nichols & ott (1978) developed a biostratigraphic scheme for species of caryapollenites and the closely related momipites for paleocene strata of the wind river basin of wyoming. however, we have found species of these two genera widely distributed in paleocene to miocene sediments, and our observations do not conform to the tight paleocene-restricted ranges proposed by nichols & ott (1978). in our experience, firm ranges for individual species of these two genera have yet to be determined, though in the labrador–baffin seaway, caryapollenites generally has an lo in the latest serravallian. caryapollenites inelegans nichols & ott 1978 (plate 18, figs 9, 10) 1978 caryapollenites inelegans nichols & ott, p. 105, 106, plate 2, figs 7, 8. age. see ‘remarks’ under the generic entry for cary a pollenites. not plotted. caryapollenites veripites (wilson & webster 1946) nichols & ott 1978 (plate 18, figs 11, 12) 1946 carya veripites wilson & webster, p. 276, fig. 14. 1978 caryapollenites veripites (wilson & webster) – nichols & ott, p. 106. age. see ‘remarks’ under the generic entry for carya pollenites. not plotted. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 82 83 genus cerebropollenites nilsson 1958 type. couper 1958, plate 30, fig. 8, as tsugaepollenites mesozoicus. 1958 cerebropollenites nilsson, p. 72. cerebropollenites mesozoicus (couper 1958) nilsson 1958 1958 tsugaepollenites mesozoicus couper, p. 155, plate 30, figs 8–10. 1958 cerebropollenites mesozoicus (couper) – nilsson, p. 72. age. lo: aptian. genus chenopodipollis krutzsch 1966 type. weyland & pflug 1957, plate 22, figs 18, 19, as periporopollenites multiplex. 1966 chenopodipollis krutzsch, p. 35. chenopodipollis sp. (plate 18, fig. 17) age. lo: miocene? genus cicatricosisporites potonié & gelletich 1933 type. potonié & gelletich 1933, plate 1, fig. 1, as cicatricosisporites dorogensis, designated by potonié (1956, p. 47). 1933 cicatricosisporites potonié & gelletich, p. 522. 1950 mohrioidites thiergart, p. 84 (name not validly published). 1951 mohrioisporites potonié, p. 144. remarks. emendations for cicatricosisporites have been proposed by potonié (1966, p. 58) and dettmann & clifford (1992, p. 289–291). cicatricosisporites minutaestriatus (bolkhovitina 1961) pocock 1964 1961 pelletieria minutaestriata bolkhovitina, p. 68, plate 20, fig. 1a–f; plate 21, fig. 3a–g. 1964 cicatricosisporites minutaestriatus (bolkhovitina) – pocock, p. 159. 1971 cicatricosisporites augustus singh, p. 68, plate 7, figs 3–11; text-fig. 7m. age. lo: early turonian. cicatricosisporites ornatus srivastava 1972 (plate 19, figs 1–4) 1972 cicatricosisporites ornatus srivastava, p. 9, plate 5, figs 3‒11; plate 6, figs 1‒4. age. lo: priabonian. remarks. the forms studied here are very similar to the type material, described by srivastava (1972) from the maastrichtian of alberta. genus cicatricososporites pflug & thomson in thomson & pflug 1953 type. selling 1944, plate 4, fig. 44, as schizaea? eocenica. 1953 cicatricososporites thomson & pflug, p. 61. 1959 schizaeoisporites krutzsch, p. 226. remarks. jansonius & hills (1976) did not agree with davies (1985) that cicatricososporites was a junior homo nym of cicatricosisporites and retained the former as a separate genus. cicatricososporites eocenicus (selling 1944) jansonius & hills 1976 (plate 18, fig. 5) 1944 schizaea? eocenica selling, p. 66, plate 4, fig. 44. 1950 sporites pseudodorogensis potonié – thiergart, p. 84; name not validly published. 1951 schizaeolsporites pseudodorogensis (potonié) – potonié, p. 144, plate 20, fig. 19; generic name not validly proposed. 1953 cicatricososporites pseudodorogensis (potonié) – thomson & pflug, p. 61. 1976 cicatricososporites eocenicus (selling) – jansonius & hills, card 468. age. lo: middle bartonian. remarks. burden & hills (1989) recorded cicatricoso sporites eocenicus from the early cretaceous, but the species is most common in the early cenozoic and especially the eocene. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 83 84 genus compositoipollenites potonié 1951 ex potonié 1960 type. potonié 1934, plate 5, fig. 25, as pollenites rhi zophorus. 1951 compositoipollenites potonié, p. 138; name not validly published. 1960 compositoipollenites potonié, p. 105. compositoipollenites sp. b of williams & brideaux 1975 (plate 18, fig. 18) 1975 compositoipollenites sp. b williams & brideaux, plate 43, fig. 15. age. lo: gelasian. remarks. compositoipollenites sp. b of williams & brideaux (1975) is very similar to the pollen of the extant genus ambrosia, being tricolporate and ornamented with short (up to about 2 μm) conate spines that are distally pointed. genus corsinipollenites nakoman 1965 type. thiergart 1940, plate 7, fig. 1 as pollenites oculus noctis. corsinipollenites oculusnoctis (thiergart 1940) nakoman 1965 (plate 18, figs 19, 20) 1940 pollenites oculus noctis thiergart, p. 47. 1965 corsinipollenites oculusnoctis (thiergart) – nakoman, p. 156. age. lo: bartonian. remarks. fossil pollen with this morphology have also been described under the modern plant name jusseia. genus extratriporopollenites pflug in thomson & pflug 1952 ex pflug in thomson & pflug 1953 type. pflug in thomson & pflug 1953, plate 6, fig. 2, as extratriporopollenites fractus. 1952 extratriporopollenites pflug in thomson & pflug, p. 14, 16; name not validly published. 1953 extratriporopollenites pflug in thomson & pflug 1952 ex pflug in thomson & pflug, p. 69. extratriporopollenites spp. (plate 19, figs 5–9) age. lo: bartonian. genus graminidites cookson 1947 ex potonié 1960 type. cookson 1947, plate 15, fig. 41, as monoporites (graminidites) media. 1947 graminidites cookson, p. 134; name not validly published. 1960 graminidites cookson ex potonié, p. 111. graminidites sp. a. of williams & brideaux 1975 (plate 19, fig. 12) 1975 graminidites spp. williams & brideaux, plate 47, figs 9, 10. age. lo: latest gelasian. genus momipites wodehouse 1933 type. momipites coryloides wodehouse 1933, fig. 43. 1933 momipites wodehouse, p. 511. remarks. see ‘remarks’ under the generic entry for caryapollenites. momipites annellus nichols & ott 1978 (plate 18, figs 13, 14) 1978 momipites annelus nichols & ott, p. 103, plate 1, figs 22–25. age. see ‘remarks’ under the generic entry for cary a pollenites. not plotted. momipites coryloides wodehouse 1933 (plate 18, fig. 15) 1933 momipites coryloides wodehouse, p. 511, fig. 43. age. see ‘remarks’ under the generic entry for cary a pollenites. not plotted. genus osmundacidites couper 1953 type. couper 1953, plate 1, fig. 5, as osmundacidites wellmanii. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 84 85 1953 osmundacidites couper, p. 20. remarks. osmundacidites is characterised by having a granulate to finely verrucate exine that is commonly irregularly organised. osmundacidites wellmannii couper 1953 (plate 20, fig. 1) 1953 osmundacidites wellmannii couper, p. 20, plate 1, fig. 5. 1959 baculatisporites wellmannii (couper) – krutzsch, p. 142. 1964 osmundacidisporites wellmannii (couper) – levet-carette, p. 98. 1968 todisporites granulatus tralau, p. 67, plate 7, fig. 1. 1972 osmundacidites araucanus volkheimer, p. 120, plate 6, figs 47‒49. 1986 osmunda sp. in williams, p. 83, plate 1, fig. 2. age. lo: burdigalian. remarks. this species has a long range, starting in the mesozoic, but its lo appears to be useful in the labrador–baffin seaway region. we have established a burdigalian lo; williams (1986) considered that osmunda sp. (= osmundacidites wellmannii) has an lo in the middle to late miocene. for a full synonymy listing of osmundacidites wellmannii, see fensome (1983, p. 323, 324). genus parviprojectus mtchedlishvili in samoilovitch & mtchedlishvili 1961 emend. braman 2013 type. samoilovitch & mtchedlishvili 1961, plate 73, fig. 2, as parviprojectus reticulatus. 1961 parviprojectus mtchedlishvili in samoilovitch & mtchedlishvili, p. 225. 2013 parviprojectus mtchedlishvili in samoilovitch & mtchedlishvili – emend. braman, p. 129. remarks. following braman (2013), parviprojectus is considered to comprise triprojectate pollen that are isopolar and have a reticulate wall. parviprojectus reticulatus mtchedlishvili in samoilovitch & mtchedlishvili 1961 (plate 17, fig. 16) 1961 parviprojectus reticulatus mtchedlishvili in samoilovitch & mtchedlishvili, p. 226, 227, plate 73, figs 2, 3. 1961 aquilapollenites reticulatus stanley, p. 348, 349, plate 8, figs 1–12. 1970 integricorpus reticulatus (mtchedlishvili) – stanley, p. 29. 1970 aquilapollenites (parviprojectus) reticulatus (mtchedlishvili) – kedves & király, p. 67. remarks. this species is characterised by its fine reticulate ornament. the above synonymy follows that of bra man (2013, p. 148), with only nomenclaturally significant entries included. age. lo: not determined; not plotted. the illustrated specimen is from late paleocene strata in rut h-11. according to braman (2013), this species ranges from the campanian to early paleocene, with a ‘provisional’ occurrence in the late eocene. the latter may be re worked, as could be the present occurrence. genus parvisaccites couper 1958 type: couper 1958, plate 29, figs 5, 6, as parvisaccites radiatus. 1958 parvisaccites couper, p. 154. parvisaccites amplus brenner 1963 1963 parvisaccites amplus brenner, p. 78, 79, plate 28, fig. 1a, b; plate 29, fig. 1a, b. age. lo: aptian. parvisaccites radiatus couper 1958 1958 parvisaccites radiatus couper, p. 154, plate 29, figs 5–8; plate 30, figs 1, 2. age. lo: aptian. genus periporopollenites pflug & thomson in thomson & pflug 1953 type. potonié 1931, plate 2, fig. 1, as pollenites stigmosus. 1953 periporopollenites pflug & thomson in thomson & pflug, p. 111. 1960 liquidambarpollenites raatz 1937 ex potonié, p. 134. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 85 86 periporopollenites sp. (plate 19, figs 10, 11) age. lo: early? miocene. not plotted. genus pistillipollenites rouse 1962 type. rouse 1962, plate 1, figs 10, 12, as pistilli polle n ites macgregorii. 1962 pistillipollenites rouse, p. 206. remarks. the distinctive characteristic of this triporate pollen grain is the pistil-like or bulbose ornamentation that covers the exine. pistillipollenites macgregorii rouse 1962 (plate 19, figs 13, 14) 1962 pistillipollenites macgregorii rouse, p. 206, plate 1, figs 8–12. age. lo: bartonian. genus quercoidites potonié et al. 1950 ex potonié 1960 type. potonié 1931, plate 2, fig. 19, as pollenites henrici. 1950 quercoidites potonié et al., p. 54; name not validly published. 1960 quercoidites potonié et al. ex potonié, p. 92. remarks. the exine of this ovoidal, tricolpate pollen grain is usually granulate to scabrate. the colpi reach almost to the poles. quercoidites sp. (plate 19, figs 15, 16) 1986 quercus form a of williams, p. 83, plate 1, fig. 8. age. lo: frequent in latest messinian. remarks. this is a finely granular form of quercoidites, also found by williams (1986) and called quercus form a. williams found that the lo of this form was in her fagus granulata zone, which she dated as early miocene. quercoidites sp. differs from quercoidites sp. a of williams & brideaux (1975) (equivalent to quer c us form b of williams 1986) in its much finer ornament. genus rugubivesiculites pierce 1961 type. pierce 1961, plate 2, fig. 57, as rugubivesiculites convolutes. 1961 rugubivesiculites pierce, p. 39. remarks. rugubivesiculites is a bivesiculate pollen with a distinctive rugulate ornamentation on the proximal surface of the central body. rugubivesiculites spp. (plate 20, fig. 2) age. lo: turonian. remarks. no attempt was made to speciate the specimens of rugubivesiculites seen in the labrador margin and offshore west greenland samples. genus tiliaepollenites potonié 1931 type. potonié 1931, fig. 14, as tiliaepollenites indu pitabilis. 1931 tiliaepollenites potonié, p. 4. 1938 tiliaepollenites raatz, p. 27. 1953 intratriporopollenites pflug & thomson in thomson & pflug, p. 87. tiliaepollenites crassipites (wodehouse 1933) comb. nov. (plate 19, figs 17, 18) 1933 tilia crassipites wodehouse, p. 515, fig. 48. 1969 tilliaepollenites crassipites (wodehouse) – penny, p. 355 (combination not validly published, basionym not fully referenced). 1975 bombacacidites sp. a williams & brideaux, plate 46, fig. 10. 1986 tilia crassipites williams, plate 2, fig. 1. age. lo: serravallian. remarks. a search of the palynodata database (palyno data inc. & white 2008) provided no indication of a previous formal transfer of this species to tiliaepollenites. tiliaepollenites sp. a (plate 19, figs 19, 20) description. this form is distinctive in its possession of a pad-like thickening at the inner end of each colpus. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 86 87 age. lo: early? miocene (not well constrained). not plotted. genus translucentipollis khlonova 1961 type. khlonova 1961, plate 16, fig. 121, as aquila pollenites plicatilis. 1961 translucentipollis khlonova, p. 89. 1966 translucentipollis khlonova – srivastava, p. 546. remarks. following braman (2013, p. 13), translucen ti pollis is considered to comprise isopolar triprojectate pollen with reduced polar projections and a more or less smooth surface. translucentipollis contiguus (tschudy 1969) braman 2013 (plate 17, fig. 13) remarks. the specimen illustrated has very fine granulae to short and delicate rugulae that are barely discernible in optical section. we consider that this wall surface fits within the granular to scabrate range of translucentipollis contiguous. age. lo: not determined. not plotted. genus wodehouseia stanley 1961 type. stanley 1961, plate 1, figs 1–3, as wodehouseia spinata. 1961 wodehouseia stanley, p. 157. remarks. wodehouseia is included in the ‘oculata’ pol len, which khlonova (1962, p. 306) defined as pollen grains with two pairs of apertures that are located close to the tips of the two long sides. wiggins (1976) gave a thorough review of various genera in the group, including wodehouseia. wodehouseia spinata stanley 1961 (plate 20, fig. 3) 1961 wodehouseia spinata stanley, p. 157, 158, 160, plate 1, figs 1–12. age. lo: latest maastrichtian. genus zlivisporis pacltová 1961 type. pacltová 1961, plate 2, figs 1–3, as zlivisporis blanensis. 1961 zlivisporis pacltová, p. 40. 1961 seductisporites khlonova, p. 56. 1962 rouseisporites pocock, p. 52. remarks. the reticulum developed on the distal surface is never fine and is often incomplete. zlivisporis spp. (plate 20, figs 4–8) age. lo: early rupelian. remarks. specimens of zlivisporis from the labrador– baffin seaway region are variable, especially in the nature of the surface reticulum, and are thus difficult to speciate. genus zonalapollenites pflug in thomson & pflug 1953 type. potonié 1931, fig. 2, as sporonites igniculus. 1934 tsugaepollenites potonié & venitz, p. 17; name not validly published. 1958 zonalapollenites pflug in thomson & pflug, p. 66. 1958 tsugaepollenites potonié & venitz ex potonié, p. 48. remarks. this genus is characterised by the numerous small, distinct equatorial sacs or vesiculae. sometimes the polar regions may be covered with greatly reduced sacs. jansonius & hills (1976, card 3265) recognised that tsugaepollenites is an obligate junior synonym of zonalapollenites. zonalapollenites igniculus (potonié 1931) thomson & pflug 1953 (plate 20, figs 9–11) 1931 sporonites igniculus potonié, p. 556, fig. 2. 1934 tsugaepollenites igniculus (potonié) – potonié & venitz, p. 17; name not validly published. 1953 zonalapollenites igniculus (potonié ex pflug) – thomson & pflug, p. 66. age. lo: latest gelasian. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 87 88 fungal elements remarks. all fungal spores are grouped together but several different morphologies are recorded, including those shown in the accompanying illustrations (plate 20, figs 12–20). fungal spores occur sporadically throughout the sections studied, but peaks were noted in the lutetian and ypresian. acknowledgements during the gestation of this paper of over a decade, we have had the unwavering support of our two institutions, the geological survey of denmark and green land (geus) and the geological survey of canada (atlantic) (gsca), part of the earth sciences sector (ess) of natural resources canada. we are also grateful for fruitful discussions with, and feedback from, kate dickie of gsca, lotte m. larsen and gunver k. pedersen of geus, stefan piasecki of the geological museum, natural history museum of denmark, university of copenhagen. we are grateful to jennifer galloway, martin pearce, jim riding and art sweet for helpful reviews that led to important improvements of the manuscript. we also extend our thanks bernie crilley and bill macmillan (gsc), and annette ryge and dorthe salomonsen (geus) for processing samples. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 88 89 agelopoulos, j. 1964: hystrichostrogylon membraniphorum n.g. n.sp. aus dem heiligenhafener kieselton (eozän). neues jahrbuch für geologie und paläontologie, monatshefte 11, 673–675. agelopoulos, j. 1967: hystrichosphären, dinoflagellaten und fora miniferen aus dem eozänen kieselton von heiligenhafen, hol stein. dissertation des grades eines dokters der mathematischnaturwissenschaftlichen fakultät der eberhard-karls-univer sität, tübingen 74 pp. (published thesis.) alberti, g. 1959: zur kenntnis der gattung deflandrea eisenack (dinoflag.) in der kreide und im alttertiär nordund mittel deutschlands. mitteilungen aus dem geologischen staatsinstitut in hamburg 28, 93–105. alberti, g. 1961: zur kenntnis mesozoischer und alttertiärer dino flagellaten und hystrichosphaerideen von nordund mittel deutschland sowie einigen anderen europäischen gebieten. palae ontographica abteilung a 116, 1–58. anderson, r.y. 1960: cretaceous–tertiary palynology, eastern side of the san juan basin, new mexico. new mexico bureau of geology and mineral resources memoir 6, 1–59. artzner, d.g. & dörhöfer, g. 1978: taxonomic note: lejeunecysta nom. nov. pro lejeunia gerlach 1961 emend. lentin and wil liams 1976 – dinoflagellate cyst genus. canadian journal of botany 56, 1381–1382. balme, b.e. 1957: spores and pollen grains from the mesozoic of western australia. commonwealth scientific and industrial research organization (australia), coal research section, ref erence t.c. 25, 1–48. barke, j., abels, h.a., sangiorgi, f., greenwood, d.r., sweet, a.r., donders, t., reichart, g-j., lotter, a.f. & brinkhuis, h. 2011: orbitally forced azolla blooms and middle eocene arctic hydrology: clues from palynology. geology 39(5), 427–430. barke, j. et al. 2012: coeval eocene blooms of the freshwater fern azolla in and around arctic and nordic seas. palaeogeography, palaeoclimatology, palaeoecology 337–338, 108–119. batten, d.j. 1996: chapter 7c. colonial chlorococcales. in: jansonius, j. & mcgregor, d.c. 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bulletin of the torrey botanical club 60, 479–524. wolff, h. 1934: mikrofossilien des pliocänen humodils der grube freigericht bei dettingen a.m. und vergleich mit älteren schich ten des tertiärs sowie posttertiären ablagerungen. preussischen geologischen landesanstalt. institut für paläobotanik und petro graphie der brennsteine arbeiten 5, 55‒86. yun hyesu 1981: dinoflagellaten aus der oberkreide (santon) von westfalen. palaeontographica abteilung b 177, 1–89. bulletin36.qxp_bulletin 36 19/12/16 13.40 side 101 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 102 103 scale the scale bar represents 20 μm in all plates. explanation in the following plates, details concerning the figured specimens are presented in a standardised, abbreviated form grouped in four categories: location locality: mainly offshore wells, but also on -• shore boreholes and outcrop sections. sample depth (in metres) in wells and boreholes;• sample number in outcrop sections, unless spe cified otherwise. sample type: cs: cuttings sample. os: outcrop section• sample. sw: sidewall core sample. cc: conventional core sample. sample number or processing number: all ‘p• numbers’ (e.g. p39553) nos are sample numbers from cana dian wells processed in canada; the remainder are ggu/geus processing numbers of both canadian and greenlandic material. slide number/letter.• optical parameters microscope identification (relating to follow-• ing vernier scale coordinates): a: leitz dialux 22 microscope 512 742/057691 at the geol ogical survey of denmark and greenland (geus). b: leica dm 2000 331596-092011 at geus. c: zeiss axioplan 2 microscope, serial no. 310243 at gsc atlantic, dartmouth, nova scotia. d: zeiss photomicroscope, serial no. 67750 at gsc atlantic, dartmouth, nova scotia. e: leitz dm rb (rs232c) at geus. vernier coordinates.• england finder coordinates.• lens magnification and type: bf: bright field.• pc: phase contrast. repositories cnlopb: slide curated in the collection of• the canada – newfoundland and labrador offshore petroleum board, st. john’s, new foundland, canada (no numbers). cnsopb: slide curated in the collection of• the canada – nova scotia offshore petroleum board, dartmouth, nova scotia, canada (no numbers). mguh: slide curated in the type collection of• the geological museum of the university of copenhagen (at the time of writing, this material is on long-term loan to geus, øster vold gade 10, dk-1350 copenhagen k, denmark). gsc: slide curated in the national collection• of type invertebrate and plant fossils, geol ogical survey of canada, 601 booth street, ot ta wa, ontario, canada k1a 0e8 (at the time of writing, this material is on long-term loan to gsc atlantic, bedford institute of oceano graphy, dartmouth, nova scotia, canada b2y 4a2). statoil: slides are stored at the offices of the• statoil petroleum company in stavanger, nor way (no numbers). all nomenclatural types are curated at either• the geological museum of the university of co pen hagen or at the geological survey of canada, as indicated above. previous publication the following figures were previously published (in black-and-white) in nøhr-hansen (2003): plate 2, fig. 12; plate 11, figs 18,19; plate 15, figs 9–12; plate 17, fig. 2. in addition, plate 6, fig. 1 was previously published in colour in nøhr-hansen (1996). plates bulletin36.qxp_bulletin 36 19/12/16 13.40 side 103 104 fig. 1. achilleodinium biformoides, optical section. location: north leif i-05, 2040‒2050 m. sample: cs, yd17597, 3. optical parameters: e, 34.6 × 17.7, k33/1, × 60 bf. repository: gsc 138151. fig. 2. achilleodinium biformoides, same specimen as fig. 1, focussed on ventral surface. location: north leif i-05, 2040‒2050 m. sample: cs, yd17597, 3. optical parameters: e, 34.6 × 17.7, k33/1, × 60 bf. repository: gsc 138152. fig. 3. achilleodinium biformoides, dorsal surface. location: south labrador n-79, 1920‒1930 m. sample: cs, p39782, 1. optical parameters: c, 192 × 1020, t31/4, × 50 bf. repository: gsc 138115. fig. 4. adnatosphaeridium vittatum, dorsal surface. location: karlsefni a-13, 1627.65–1636.80 m. sample: cs, p39577, 1. optical parameters: d, 179 × 866, r29/3, × 40 pc. repository: gsc 138013. fig. 5. alisocysta margarita, dorsal surface. location: north leif i-05, 2310‒2320 m. sample: cs, yd17606, 2. optical parameters: e, 28.5 × 6.9, u27/3, × 60 bf. repository: mguh 31266. fig. 6. alisocysta circumtabulata, ventral view of ventral surface. location: south labrador n-79, 3060‒3070 m. sample: cs, p39819, 1. optical parameters: c, 166 × 1026, r32/0–2, × 50 bf. repository: gsc 138124. fig. 7. alisocysta circumtabulata, same specimen as fig. 6, focussed on dorsal surface. location: south labrador n-79, 3060‒3070 m. sample: cs, p39819, 1. optical parameters: c, 166 × 1026, r32/0–2, × 50 bf. repository: gsc 138124 fig. 8. adnatosphaeridium vittatum, apical view. location: south labrador n-79, 2310‒2320 m. sample: cs, p39795, 1. optical parameters: c, 184 × 1028, s32/4, × 50 bf. repository: gsc 138120. fig. 9. alisocysta margarita, dorsal view. location: north leif i-05, 2370‒2380 m. sample: cs, yd17608, 4. optical parameters: e, 38.9 × 13.6, o37/2, × 60 bf. repository: gsc 138161. fig. 10. alterbidinium acutulum, dorsal view of dorsal surface. location: karlsefni a-13, 1106.44–1115.58 m. sample: cs, p39558, 1. optical parameters: d, 192 × 918, t34/2, × 40 pc. repository: gsc 138010. fig. 11. alterbidinium acutulum, dorsal surface. location: skolp e-07, 1295 m. sample: cs, yd15594, 3. optical parameters: a, 49.0 × 113.2, c50/1, × 60 bf. repository: mguh 31267. fig. 12. alterbidinium biaperturum, dorsal view of dorsal surface. location: skolp e-07, 1070 m. sample: cs, yd15580, 3. optical parameters: a, 19.6 × 96.4, u19/2, × 60 bf. repository: mguh 31268. fig. 13. alterbidinium biaperturum, same specimen as fig. 12, dorsal view of ventral surface. location: skolp e-07, 1070 m. sample: cs, yd15580, 3. optical parameters: a, 19.6 × 96.4, u19/2, × 60 bf. repository: mguh 31268. fig. 14. alterbidinium biaperturum, dorsal view of ventral surface. location: skolp e-07, 1250 m. sample: cs, yd15592, 3. optical parameters: a, 25.2 × 98.0, s25/3, × 60 bf. repository: mguh 31269. fig. 15. alterbidinium? bicellulum, ventral view. location: gjoa o-37, 1620 m. sample: cs, yd16074, 2. optical parameters: a, 16.2 × 100.0, r16/1, × 60 bf. repository: mguh 31270. fig. 16. alterbidinium? bicellulum, ventral view. location: gjoa o-37, 1620 m. sample: cs, yd16074, 2. optical parameters: a, 29.2 × 93.0, y29/0, × 60 bf. repository: mguh 31271. fig. 17. alterbidinium ioannidesii, dorsal view of ventral surface. location: skolp e-07, 2390 m. sample: cs, yd15666, 2. optical parameters: a, 20.1 × 113.2, b20/3, × 60 bf. repository: mguh 31272. fig. 18. alterbidinium varium, ventral view. location: skolp e-07, 1715 m. sample: cs, jeh15622, 3. optical parameters: a, 40.0 × 103.7, m40/4, × 60 bf. repository: mguh 31273. fig. 19. alterbidinium varium, dorsal view. location: skolp e-07, 1715 m. sample: cs, jeh15622, 5. optical parameters: a, 26.0 × 101.6, p26/3, × 60 bf. repository: cnlopb. fig. 20. apectodinium homomorphum, dorsal view. location: north leif i-05, 2010‒2020 m. sample: cs, yd17596, 2. optical parameters: e, 34.1 × 12.8, o33/3, × 60 bf. repository: mguh 31274. plate 1 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 104 105 1 9 5 13 17 2 10 6 14 18 3 11 7 15 19 4 12 8 16 20 plate 1 plate 1 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 105 106 fig. 1. apectodinium parvum, ventral view. location: north leif i-05, 2100‒2110 m. sample: cs, yd17599, 3. optical parameters: e, 22.6× 16.8, k20/4, × 60 bf. repository: gsc 138154. fig. 2. apectodinium parvum, dorsal view. location: north leif i-05, 2070‒2080 m. sample: cs, yd17598, 4. optical parameters: e, 20.3 × 23.3, d18/0, × 60 bf. repository: mguh 31275. fig. 3. apectodinium parvum, dorsal view of dorsal surface. location: rut h-11, 3335‒3345 m. sample: cs, p39405, 1. optical parameters: d, 121 × 1034, m46/2, × 40 pc. repository: gsc 137967. fig. 4. apectodinium quinquelatum, dorsal view of ventral surface. location: north leif i-05, 2250‒2260 m. sample: cs, yd17604, 3. optical parameters: e, 55.2 × 18.9, g54/4, × 60 bf. repository: mguh 31276. fig. 5. aptea polymorpha, dorsal view. location: roberval k-92, 3160‒3170 m. sample: cs, p2008177, 1. optical parameters: c, 79 × 980, h27/7, × 50 bf. repository: gsc 137910. fig. 6. apteodinium australiense, dorso-ventral view. location: karlsefni a-13, 969.28–978.42 m. sample: cs, p39553, 1. optical parameters: d, 167 × 898, q32/4, × 40 pc. repository: gsc 138005. fig. 7. apteodinium spiridoides, right lateral view. location: karlsefni a-13, 2505.49–2514.63 m. sample: cs, p39608, 1. optical parameters: d, 171 × 990, r42/0, × 40 pc. repository: gsc 138037. fig. 8. apteodinium spiridoides, right lateral view. location: rut h-11, 725‒735 m. sample: cs, p39318, 1. optical parameters: d, 78 × 913, g34/0‒3, × 40 pc. repository: gsc 137919. fig. 9. areoligera circumsenonensis, ventral view of dorsal surface. location: karlsefni a-13, 2615.22–2624.36 m. sample: cs, p39612, 1. optical parameters: d, 183 × 990, s42/1‒3, × 40 pc. repository: gsc 138040. fig. 10. areoligera gippingensis, dorso-ventral view. location: karlsefni a-13, 3822.24–3831.38 m. sample: cs, p39655, 1. optical parameters: d, 143 × 930, o36/1, × 40 pc. repository: gsc 138054. fig. 11. areoligera gippingensis, dorso-ventral view. location: gjoa o-37, 2240 m. sample: cs, yd16095, 2. optical parameters: e, 52.3 × 14.2, m51/4, × 60 bf. repository: mguh 31277. fig. 12. areosphaeridium diktyoplokum, dorso-ventral view. location: hellefisk-1, 723 m. sample: sw, yd13709, 2. optical parameters: a, 38.9 × 108.4, h39/4, × 60 bf. repository: mguh 26473. fig. 13. atopodinium cf. haromense, dorso-ventral view. location: skolp e-07, 1550 m. sample: cs, jeh15611, 5. optical parameters: a, 35.0 × 95.5, v35/1, × 60 bf. repository: mguh 31278. fig. 14. axiodinium augustum, dorso-ventral view. location: north leif i-05, 2250‒2260 m. sample: cs, yd17604, 3. optical parameters: e, 43.0 × 21.6, e42/1, × 60 bf. repository: mguh 31279. fig. 15. axiodinium augustum, dorso-ventral view, same specimen as fig. 14, focussed on antapical part. location: north leif i-05, 2250‒2260 m. sample: cs, yd17604, 3. optical parameters: e, 43.0 × 21.6, e42/1, × 60 bf. repository: mguh 31279. fig. 16. batiacasphaera micropapillata, dorso-ventral view. location: rut h-11, 845‒855 m. sample: cs, p39322, 1. optical parameters: d, 93 × 1020, j45/0, × 40 pc. repository: gsc 137924. fig. 17. batioladinium jaegeri, ventral view of ventral surface. location: skolp e-07, 1655 m. sample: cs, jeh15618, 3. optical parameters: a, 45.7 × 107.9, h46/4, × 60 bf. repository: mguh 31281. fig. 18. callaiosphaeridium asymmetricum, apical-antapical view. location: skolp e-07, 1535 m. sample: cs, jeh15610, 3. optical parameters: a, 33.7 × 98.8, r34/3, × 60 bf. repository: mguh 31282. fig. 19. cerebrocysta bartonensis, left lateral view. location: ogmund e-72, 1336 m. sample: sw, jeh15552, 2. optical parameters: a, 39.8 × 107.7, h40/4, × 60 bf. repository: mguh 31283. fig. 20. cerebrocysta bartonensis, left lateral view, same specimen as fig. 19, focussed on operculum, within cyst. location: ogmund e-72, 1336 m. sample: sw, jeh15552, 2. optical parameters: a, 39.8 × 107.7, h40/4, × 60 bf. repository: mguh 31283. plate 2 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 106 107 1 9 13 17 2 10 14 18 3 11 7 15 4 12 8 16 20 plate 2 65 19 plate 2 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 107 108 fig. 1. cerodinium diebelii, dorso-ventral view. location: roberval k-92, 3160‒3170 m. sample: cs, p2008177, 1. optical parameters: c, 130 × 918, n20/0, × 40 bf. repository: gsc 137909. fig. 2. cerodinium diebelii, ventral view of dorsal surface. location: roberval k-92, 2500‒2510 m. sample: cs, p19229, 1. optical parameters: c, 100 × 1018, k31/0–2, × 50 bf. repository: gsc 137905. fig. 3. cerodinium diebelii, dorsal view of dorsal surface. location: north leif i-05, 2370‒2380 m. sample: cs, yd17608, 2. optical parameters: e, 39.6 × 16.1, l38/2, × 60 bf. repository: mguh 31284. fig. 4. cerodinium diebelii, dorso-ventral view. location: north leif i-05, 2520‒2530 m. sample: cs, yd17613, 3. optical parameters: e, 23.8 × 18.7, h22/3, × 60 bf. repository: mguh 31285. fig. 5. cannosphaeropsis passio, oblique antapical view. location: qulleq-1, 1847 m. sample: sw, yd14581, 4. optical parameters: a, 38.4 × 108.0, j38/2, × 40 bf. repository: statoil. fig. 6. cerebrocysta magna, oblique right lateral view. location: qulleq-1, 1862.1 m. sample: sw, yd14587, 3. optical parameters: a, 29.8 × 100.7, q30/3, × 60 bf. repository: statoil/mguh 26505. fig. 7. cerodinium glabrum, dorsal surface. location: gilbert f-53, 3420‒3430 m. sample: cs, p17040, 1. optical parameters: c, 158 × 1010, l30/2, × 50 bf. repository: gsc 137887. fig. 8. cerodinium glabrum, dorsal surface. location: gjoa o-37, 2400 m. sample: cs, yd16097, 2. optical parameters: e, 21.4 × 25.3, b19/2, × 60 bf. repository: mguh 31286. fig. 9. cerodinium kangiliense, dorsal view of dorsal surface. location: skolp e-07, 1460 m. sample: cs, yd15605, 3. optical parameters: a, 52.0 × 102.8, n53/3, × 60 bf. repository: mguh 31287. fig. 10. cerodinium kangiliense, dorsal view, same specimen as fig. 9, focussed on ventral surface. location: skolp e-07, 1460 m. sample: cs, yd15605, 3. optical parameters: a, 52.0 × 102.8, n53/3, × 60 bf. repository: mguh 31287. fig. 11. cerodinium striatum, dorso-ventral view of dorsal surface. location: gilbert f-53, 2410‒2420 m. sample: cs, p39493, 1. optical parameters: c, 70 × 1040, g34/1, × 50 bf. repository: gsc 137983. fig. 12. cerodinium striatum, dorso-ventral view. location: rut h-11, 2765‒2775 m. sample: cs, p39386, 1. optical parameters: d, 65 × 869, f29/2, × 40 pc. repository: gsc 137955. fig. 13. cerodinium speciosum, dorso-ventral view of dorsal surface. location: hekja o-71, 4510‒4520 m. sample: cs, p20645, 1. optical parameters: c, 50 × 1007, e30/0, × 50 bf. repository: gsc 137915. fig. 14. chatangiella madura, dorsal view of dorsal surface. location: skolp e-07, 1580 m. sample: cs, jeh15613, 3. optical parameters: a, 31.0 × 112.8, c31/4, × 60 bf. repository: mguh 31288. fig. 15. chatangiella madura, dorsal view, same specimen as fig. 14, focussed on optical section. location: skolp e-07, 1580 m. sample: cs, jeh15613, 3. optical parameters: a, 31.0 × 112.8, c31/4, × 60 bf. repository: mguh 31288. fig. 16. chatangiella tripartita, dorsal view of dorsal surface. location: south labrador n-79, 1230‒1240 m. sample: cs, p39761, 1. optical parameters: c, 178 × 988, s28/0–2, × 50 bf. repository: gsc 138081. fig. 17. chiropteridium galea, dorsal view. location: karlsefni a-13, 1051.57–1060.72 m. sample: cs, p39556, 1. optical parameters: d, 228 × 1029, x46/1, × 40 pc. repository: gsc 138008. fig. 18. chiropteridium galea, ventral view of ventral surface. location: karlsefni a-13, 1079.01–1088.15 m. sample: cs, p39557, 1. optical parameters: d, 170 × 991, r42/1, × 40 pc. repository: gsc 138009. fig. 19. chiropteridium galea, ventral view. location: karlsefni a-13, 1161.30–1170.45 m. sample: cs, p39560, 1. optical parameters: d, 172 × 889, r31/2, × 40 pc. repository: gsc 138011. fig. 20. chatangiella tripartita, ventral view of dorsal surface. location: karlsefni a-13, 2267.74–2276.88 m. sample: cs, p39599, 1. optical parameters: d, 192 × 926, t35/0, × 40 pc. repository: gsc 138029. plate 3 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 108 109 1 9 5 13 17 2 10 6 14 18 3 11 7 15 19 4 12 8 16 20 plate 3 plate 3 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 109 110 fig. 1. chiropteridium gilbertii sp. nov., dorso-ventral view. location: south labrador n-79, 1440‒1450 m. sample: cs, p39768, 1. optical parameters: c, 56 × 1028, e32/4, × 50 bf. repository: gsc 138112. fig. 2. chiropteridium gilbertii sp. nov., ventral view of dorsal surface. location: south labrador n-79, 1560‒1570 m. sample: cs, p39772, 1. optical parameters: c, 168 × 1031, r33/1, × 50 bf. repository: gsc 138113. fig. 3. chiropteridium gilbertii sp. nov., ventral view, same specimen as fig. 2, focussed on ventral surface. location: south labrador n-79, 1560‒1570 m. sample: cs, p39772, 1. optical parameters: c, 168 × 1031, r33/1, × 50 bf. repository: gsc 138113. fig. 4. chiropteridium gilbertii sp. nov., holotype, ventral view. location: gilbert f-53, 1600‒1610 m. sample: cs, p39466, 1. optical parameters: c, 174 × 1060, r36/3, × 50 bf. repository: gsc 137976. fig. 5. chlamydophorella nyei, left lateral view. location: south coast bylot island, section b, 1.8 m above base (sample hfb09-17a). sample: os, p5148-4, c. optical parameters: c, 193 × 901, h57/0–1, × 50 bf. repository: gsc 138130. fig. 6. chlamydophorella cf. nyei, dorso-ventral view. location: ogmund e-72, 1620 m. sample: cs, yd15740, 5. optical parameters: a, 19.5 × 110.5, f19/2, × 60 bf. repository: gsc 138145. fig. 7. chlamydophorella cf. nyei, right lateral view. location: north leif i-05, 2670–2680 m. sample: cs, yd17618, 2. optical parameters: e, 18.6 × 19.5, g16/2, × 60 bf. repository: mguh 31289. fig. 8. chlamydophorella cf. nyei, dorso-ventral view. location: skolp e-07, 1685 m. sample: cs, jeh15620, 3. optical parameters: b, 30.47 × 5.36, x32/1, × 60 bf. repository: mguh 31290. fig. 9. charlesdowniea coleothrypta, ventral view of ventral surface. location: gilbert f-53, 2050‒2060 m. sample: cs, p39481, 1. optical parameters: c, 57 × 1014, e31/3, × 50 bf. repository: gsc 137981. fig. 10. cleistosphaeridium elegantulum sp. nov., apical view. location: karlsefni a-13, 2203.73–2212.87 m. sample: cs, p39597, 1. optical parameters: d, 122 × 866, m29/0‒1, × 40 pc. repository: gsc 138026. fig. 11. cleistosphaeridium elegantulum sp. nov., oblique apical view. location: karlsefni a-13, 2203.73–2212.87 m. sample: cs, p39597, 1. optical parameters: d, 230 × 894, x32/0, × 40 pc. repository: gsc 138027. fig. 12. cleistosphaeridium elegantulum sp. nov., holotype, dorsoventral view. location: karlsefni a-13, 2286.03–2295.17 m. sample: cs, p39600, 1. optical parameters: d, 179 × 1017, s44/2, × 40 pc. repository: gsc 138033. fig. 13. cleistosphaeridium palmatum sp. nov., holotype, dorsal view of optical section. location: roberval k-92, 2450‒2460 m. sample: cs, p17706, 1. optical parameters: c, 69 × 1006, g30/0–2, × 50 bf. repository: gsc 137897. fig. 14. cleistosphaeridium palmatum sp. nov., holotype, dorsal view, focussed on dorsal surface. location: roberval k-92, 2450‒2460 m. sample: cs, p17706, 1. optical parameters: c, 69 × 1006, g30/0–2, × 50 bf. repository: gsc 137897. fig. 15. cleistosphaeridium palmatum sp. nov., dorso-ventral view. location: rut h-11, 3005‒3015 m. sample: cs, p39394, 1. optical parameters: d, 141 × 905, o33/1, × 40 pc. repository: gsc 137960. fig. 16. cleistosphaeridium palmatum sp. nov., dorso-ventral view. location: rut h-11, 3275‒3285 m. sample: cs, p39403, 1. optical parameters: d, 140 × 925, n35/3‒4, × 40 pc. repository: gsc 137966. fig. 17. chytroeisphaeridia hadra sp. nov., dorsal surface showing the archaeopyle. location: roberval k-92, 1700‒1710 m. sample: cs, p17681, 1. optical parameters: c, 48 × 1017, e30/2, × 50 bf. repository: gsc 137888. fig. 18. chytroeisphaeridia hadra, sp. nov., holotype, optical section view. location: roberval k-92, 3120‒3140 m. sample: cs, p17728, 1. optical parameters: c, 195 × 1098, u40/1, × 50 bf. repository: gsc 137902. fig. 19. cleistosphaeridium diversispinosum, dorso-ventral view. location: karlsefni a-13, 969.28–978.42 m. sample: cs, p39553, 1. optical parameters: d, 163 × 971, q40/0, × 40 pc. repository: gsc 138004. fig. 20. cleistosphaeridium diversispinosum, operculum. location: karlsefni a-13, 2203.73–2212.87 m. sample: cs, p39597, 1. optical parameters: d, 115 × 972, l40/0, × 40 pc. repository: gsc 138025. plate 4 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 110 111 1 9 5 13 17 2 10 6 14 18 3 11 7 15 19 4 12 8 16 20 plate 4 plate 4 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 111 112 fig. 1. cleistosphaeridium polypetellum, dorso-ventral view. location: karlsefni a-13, 2944.40–2953.55 m. sample: cs, p39624, 1. optical parameters: d, 215 × 1017, v44/4, × 40 pc. repository: gsc 138047. fig. 2. cleistosphaeridium polypetellum. location: karlsefni a-13, 3054.13–3063.28 m. sample: cs, p39628, 1. optical parameters: d, 151 × 896, p32/1‒2, × 40 pc. repository: gsc 138048. fig. 3. cordosphaeridium delimurum, dorso-ventral view. location: karlsefni a-13, 2807.24–2816.39 m. sample: cs, p39619, 1. optical parameters: d, 166 × 976, q40/0‒4, × 40 pc. repository: gsc 138046. fig. 4. cordosphaeridium delimurum, dorso-ventral view, same specimen as fig. 3. location: karlsefni a-13, 2807.24–2816.39 m. sample: cs, p39619, 1. optical parameters: d, 166 × 976, q40/0‒4, × 40 pc. repository: gsc 138046. fig. 5. cordosphaeridium fibrospinosum, oblique apical view. location: karlsefni a-13, 3989.88–3999.02 m. sample: cs, p39661, 1. optical parameters: d, 122 × 980, m41/0‒1, × 40 pc. repository: gsc 138060. fig. 6. cordosphaeridium fibrospinosum, dorsal view. location: karlsefni a-13, 3989.88–3999.02 m. sample: cs, p39661, 1. optical parameters: d, 126 × 897, m32/0, × 40 pc. repository: gsc 138061. fig. 7. cordosphaeridium cantharellus, right lateral view. location: rut h-11, 755‒765 m. sample: cs, p39319, 1. optical parameters: d, 110 × 978, l40/2, × 40 pc. repository: gsc 137922. fig. 8. cordosphaeridium funiculatum, left lateral view. location: ralegh n-18, 1525 m. sample: cs, yd16219, 2. optical parameters: e, 43.3 × 23.8, c42/1, × 60 bf. repository: mguh 31291. fig. 9. cordosphaeridium gracile, dorsal surface. location: karlsefni a-13, 3081.57–3090.71 m. sample: cs, p39629, 1. optical parameters: d, 171 × 1030, r46/0, × 40 pc. repository: gsc 138049. fig. 10. cordosphaeridium inodes, dorsal surface. location: karlsefni a-13, 3931.97–3941.11 m. sample: cs, p39659, 1. optical parameters: d, 97 × 910, j34/1‒3, × 40 pc. repository: gsc 138059. fig. 11. dapsilidinium pseudoinsertum sp. nov., dorso-ventral view. location: karlsefni a-13, 2697.51–2706.66 m. sample: cs, p39615, 1. optical parameters: d, 98 × 946, j37/4, × 40 pc. repository: gsc 138043. fig. 12. dapsilidinium pseudoinsertum sp. nov., holotype, dorsoventral view. location: rut h-11, 2825‒2835 m. sample: cs, p39388, 1. optical parameters: d, 196 × 1051, u48/1‒2, × 40 pc. repository: gsc 137957. fig. 13. dapsilidinium pseudocolligerum. location: gilbert f-53, 2020‒2030 m. sample: cs, p39480, 1. optical parameters: c, 45 × 1008, d30/0–4, × 50 bf. repository: gsc 137980. fig. 14. dapsilidinium simplex, dorso-ventral view. location: rut h-11, 3845‒3855 m. sample: cs, p39421, 1. optical parameters: d, 165 × 921, q35/3, × 40 pc. repository: gsc 137974. fig. 15. deflandrea denticulata, dorso-ventral view of dorsal surface. location: karlsefni a-13, 4038.65–4047.79 m. sample: cs, p39663, 1. optical parameters: d, 165 × 1014, q44/0, × 40 pc. repository: gsc 138065. fig. 16. deflandrea denticulata, dorso-ventral view of ventral surface. location: karlsefni a-13, 4038.65–4047.79 m. sample: cs, p39663, 1. optical parameters: d, 60 × 1068, e50/3, × 40 pc. repository: gsc 138064. fig. 17. deflandrea borealis sp. nov., dorso-ventral view. location: karlsefni a-13, 2697.51–2706.66 m. sample: cs, p39615, 1. optical parameters: d, 169 × 929, q35/4, × 40 pc. repository: gsc 138044. fig. 18. deflandrea borealis sp. nov., dorso-ventral view. location: rut h-11, 2915‒2925 m. sample: cs, p39391, 1. optical parameters: d, 198 × 883, u31/1, × 40 pc. repository: gsc 137959. fig. 19. deflandrea borealis sp. nov., holotype, dorso-ventral view. location: rut h-11, 3095‒3105 m. sample: cs, p39397, 1. optical parameters: d, 189 × 986, t41/2, × 40 pc. repository: gsc 137964. fig. 20. deflandrea borealis sp. nov., dorso-ventral view. location: gilbert f-53, 1810‒1820 m. sample: cs, p39473, 1. optical parameters: c, 175 × 1014, r31/3, × 50 bf. repository: gsc 137977. plate 5 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 112 113 1 9 5 13 17 2 10 6 14 18 3 11 7 15 19 4 12 8 16 20 plate 5 plate 5 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 113 114 fig. 1. deflandrea galeata, dorso-ventral view of dorsal surface. location: annertuneq, 388 m, ggu366591. sample: os, c402-g, 5. optical parameters: a, 56.1 × 112.8, b556/4, × 60 bf. repository: mguh 23924. fig. 2. deflandrea majae, ventral view of ventral surface. location: skolp e-07, 1055 m. sample: cs, yd15579, 4. optical parameters: a, 45.7 × 96.9, t46/4, × 60 bf. repository: gsc 138142. fig. 3. deflandrea oebisfeldensis, dorso-ventral view of ventral surface. location: hekja o-71, 3360 m. sample: cs, jeh16039, 3. optical parameters: e, 32.1 × 5.6, w31/1, × 60 bf. repository: mguh 31292. fig. 4. deflandrea phosphoritica, dorso-ventral view. location: snorri j-90, 2249.45‒2256.6 m. sample: cs, p9747, 10. optical parameters: c, 220 × 1035, w33/0, × 50 bf. repository: gsc 138136. fig. 5. diphyes brevispinum, dorso-ventral view. location: south labrador n-79, 1920‒1930 m. sample: cs, p39782, 1. optical parameters: c, 118 × 1030, m32/2–m33/1, × 50 bf. repository: gsc 138114. fig. 6. diphyes brevispinum, dorso-ventral view. location: south labrador n-79, 1980‒1990 m. sample: cs, p39784, 1. optical parameters: c, 144 × 1042, o34/3, × 50 bf. repository: gsc 138117. fig. 7. diphyes ficusoides, dorso-ventral view. location: bjarni o-82, 1905‒1915 m. sample: cs, p39718, 1. optical parameters: c, 130 × 1061, n36/1–2, × 50 bf. repository: gsc 138071. fig. 8. diphyes ficusoides, dorso-ventral view. location: rut h-11, 3125‒3135 m. sample: cs, p39398, 1. optical parameters: d, 164 × 971, q40/0, × 40 pc. repository: gsc 137965. fig. 9. disphaerogena carposphaeropsis. location: annertuneq, 451 m, ggu405093. sample: os, yd11738, 3. optical parameters: a, 27.6 × 96.7, u27/2, × 40 bf. repository: mguh 31293. fig. 10. cyclonephelium distinctum, dorsal view of dorsal surface. location: gilbert f-53, 3220‒3230 m. sample: cs, p39520, 1. optical parameters: c, 28 × 1026, b32/3–4, × 50 bf. repository: gsc 137990. fig. 11. dinogymnium longicorne. location: skolp e-07, 1895 m. sample: cs, jeh15633, 4. optical parameters: a, 42.9 × 110.5, e43/4, × 60 bf. repository: mguh 31294. fig. 12. diphyes colligerum, dorso-ventral view. location: gilbert f-53, 1990‒2000 m. sample: cs, p39479, 1. optical parameters: c, 194 × 1058, t35/4, × 50 bf. repository: gsc 137979. fig. 13. eatonicysta furensis, dorsal view of dorsal surface. location: north leif i-05, 2070‒2080 m. sample: cs, yd17598, 3. optical parameters: e, 19.3 × 15.7, m17/0, × 60 bf. repository: gsc 138153. fig. 14. eatonicysta ursulae, dorso-ventral view. location: south labrador n-79, 2220‒2230 m. sample: cs, p39792, 1. optical parameters: c, 56 × 1042, e34/3, × 50 bf. repository: gsc 138119. fig. 15. evittosphaerula? foraminosa sp. nov., apical view. location: north leif i-05, 2130‒2140 m. sample: cs, yd17600, 3. optical parameters: e, 23.7 × 10.0, s22/1, × 60 bf. repository: gsc 138155. fig. 16. evittosphaerula? foraminosa sp. nov., apical view of apical surface. location: north leif i-05, 2190‒2200 m. sample: cs, yd17602, 3. optical parameters: e, 36.5 × 26.0, a35/0, × 60 bf. repository: gsc 138160. fig. 17. evittosphaerula? foraminosa sp. nov., oblique apical view of epicyst. location: north leif i-05, 2130‒2140 m. sample: cs, yd17600, 3. optical parameters: e, 32.8 × 20.8, f31/3, × 60 bf. repository: gsc 138157. fig. 18. evittosphaerula? foraminosa sp. nov., oblique antapical view. location: north leif i-05, 2130‒2140 m. sample: cs, yd17600, 3. optical parameters: e, 34.2 × 5.7, w32/2, × 60 bf. repository: gsc 138158. fig. 19. evittosphaerula? foraminosa sp. nov., holotype, right late ral view focussed on right lateral surface. location: north leif i-05, 2130‒2140 m. sample: cs, yd17600, 3. optical parameters: e, 44.9 × 15.5, m43/0, × 60 bf. repository: gsc 138159. fig. 20. evittosphaerula? foraminosa sp. nov., holotype, right late ral view, focussed on left lateral surface. location: north leif i-05, 2130‒2140 m. sample: cs, yd17600, 3. optical parameters: e, 44.9 × 15.5, m43/0, × 60 bf. repository: gsc 138159. plate 6 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 114 115 1 9 5 13 17 2 10 6 14 18 3 11 7 15 19 4 12 8 16 20 plate 6 plate 6 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 115 116 fig. 1. enneadocysta magna, ventral view. location: rut h-11, 1205‒1215 m. sample: cs, p39334, 1. optical parameters: d, 100 × 990, j42/3, × 40 pc. repository: gsc 137929. fig. 2. enneadocysta magna, ventral view. location: rut h-11, 1235‒1245 m. sample: cs, p39335, 1. optical parameters: d, 216 × 859, v28/3‒4, × 40 pc. repository: gsc 137933. fig. 3. eocladopyxis peniculata. location: south labrador n-79, 2310‒2320 m. sample: cs, p20171, 1. optical parameters: c, 190 × 1100, t41/0–3, × 50 bf. repository: gsc 137911. fig. 4. fibrocysta bipolaris, right lateral view. location: roberval k-92, 2390‒2400 m. sample: cs, p17704, 1. optical parameters: c, 33 × 1090, c39/0, × 50 bf. repository: gsc 137895. fig. 5. ginginodinium? flexidentatum sp. nov., dorso-ventral view of ventral surface. location: north leif i-05, 2010‒2020 m. sample: cs, yd17596, 3. optical parameters: e, 24.9 × 21.8, e23/3, × 60 bf. repository: gsc 138150. fig. 6. ginginodinium? flexidentatum sp. nov., dorso-ventral view of dorsal surface. location: ogmund e-72, 1340 m. sample: cs, yd15727, 3. optical parameters: a, 43.4 × 96.5, u44/1, × 60 bf. repository: mguh 31295. fig. 7. ginginodinium? flexidentatum sp. nov., dorso-ventral view of ventral surface. location: hekja o-71, 3090 m. sample: cs, jeh16029, 2. optical parameters: e, 35.6 × 20.8, f34/0, × 60 bf. repository: mguh 31296. fig. 8. ginginodinium? flexidentatum sp. nov., dorso-ventral view. location: karlsefni a-13, 2807.24–2816.39 m. sample: cs, p39619, 1. optical parameters: d, 146 × 1060, o49/0, × 40 pc. repository: gsc 138045. fig. 9. ginginodinium? flexidentatum sp. nov., dorso-ventral view. location: north leif i-05, 2010‒2020 m. sample: cs, yd17596, 3. optical parameters: a, 22.5 × 19.3, h20/0, × 60 bf. repository: gsc 138149. fig. 10. ginginodinium? flexidentatum sp. nov., dorso-ventral view. location: rut h-11, 2915‒2925 m. sample: cs, p39391, 1. optical parameters: d, 140 × 1029, o46/1, × 40 pc. repository: gsc 137958. fig. 11. ginginodinium? flexidentatum sp. nov., holotype, dorsoventral view. location: bjarni o-82, 1815‒1825 m. sample: cs, p39715, 1. optical parameters: c, 36 × 1028, c32/4, × 50 bf. repository: gsc 138070. fig. 12. gillinia hymenophora. location: gilbert f-53, 3280‒3290 m. sample: cs, p39522, 1. optical parameters: c, 155 × 1045, p34/3–4, × 50 bf. repository: gsc 137991. fig. 13. habibacysta tectata, right lateral view, showing detached operculum. location: karlsefni a-13, 667.52–676.66 m. sample: cs, p39542, 1. optical parameters: d, 180 × 979, s41/1, × 40 pc. repository: gsc 137995. fig. 14. glaphyrocysta exuberans, dorso-ventral view. location: north leif i-05, 1530‒1540 m. sample: cs, yd15580, 2. optical parameters: a, 38.8 × 23.5, c37/4, × 60 bf. repository: mguh 31297. fig. 15. glaphyrocysta divaricata, dorso-ventral view. location: karlsefni a-13, 3438.19–3447.33 m. sample: cs, p39642, 1. optical parameters: d, 194 × 1048, t48/3, × 40 pc. repository: gsc 138053. fig. 16. glaphyrocysta divaricata, dorso-ventral view. location: karlsefni a-13, 3438.19–3447.33 m. sample: cs, p39642, 1. optical parameters: d, 188 × 1032, s46/3‒4, × 40 pc. repository: gsc 138052. fig. 17. glaphyrocysta retiintexta, ventral view. location: karlsefni a-13, 4038.65–4047.79 m. sample: cs, p39663, 1. optical parameters: d, 220 × 1066, w49/2‒4, × 40 pc. repository: gsc 138066. fig. 18. histiocysta palla, ventral view. location: ogmund e-72, 1620 m. sample: cs, yd15740, 4. optical parameters: a, 37.0 × 98.7, r37/4, × 60 bf. repository: mguh 31298. fig. 19. histiocysta palla, oblique right lateral view, same specimen as fig. 18. location: ogmund e-72, 1620 m. sample: cs, yd15740, 4. optical parameters: a, 37.0 × 98.7, r37/4, × 60 bf. repository: mguh 31298. fig. 20. histiocysta palla, dorsal surface. location: bjarni o-82, 2300‒2310 m. sample: cs, p39731, 1. optical parameters: c, 140 × 996, o29/0, × 50 bf. repository: gsc 138073. plate 7 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 116 117 1 9 5 13 17 2 10 6 14 18 3 11 7 15 19 4 12 8 16 20 plate 7 plate 7 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 117 118 fig. 1. glaphyrocysta texta, oblique lateral view. location: north leif i-05, 2100‒2110 m. sample: cs, yd17599, 4. optical parameters: e, 25.0 × 16.6, l23/2, × 60 bf. repository: mguh 31299. fig. 2. glaphyrocysta vicina, ventral view of ventral surface. location: ralegh n-18, 1725 m. sample: cs, yd16224, 2. optical parameters: e, 45.9 × 7.7, t45/3, × 60 bf. repository: mguh 31300. fig. 3. heterosphaeridium bellii. location: gilbert f-53, 3460‒3470 m. sample: cs, p39528, 1. optical parameters: c, 62 × 1062, f36/0, × 50 bf. repository: gsc 137992. fig. 4. heterosphaeridium bellii. location: north leif i-05, 2670‒2680 m. sample: cs, yd17618, 3. optical parameters: e, 43.0 × 7.1, v41/2, × 60 bf. repository: gsc 138166. fig. 5. homotryblium tenuispinosum, apical view, focussed on archaeopyle margin. location: gilbert f-53, 1960‒1970 m. sample: cs, p39478, 1. optical parameters: c, 194 × 983, t28/3, × 50 bf. repository: gsc 137978. fig. 6. homotryblium tenuispinosum, apical view, same specimen as fig. 5, focussed on hypocyst. location: gilbert f-53, 1960‒1970 m. sample: cs, p39478, 1. optical parameters: c, 194 × 983, t28/3, × 50 bf. repository: gsc 137978. fig. 7. homotryblium tenuispinosum, dorso-ventral view. location: south labrador n-79, 1920‒1930 m. sample: cs, p39782, 1. optical parameters: c, 40 × 1061, d35/1, × 50 bf. repository: gsc 138116. fig. 8. heterosphaeridium difficile, dorso-ventral view. location: south labrador n-79, 3485‒3495 m. sample: cs, p20210, 1. optical parameters: c, 168 × 1128, r43/1‒2, × 50 bf. repository: gsc 137912. fig. 9. homotryblium abbreviatum, dorsal view. location: snorri j-90, 2249.45‒2256.6 m. sample: cs, p9747, 10. optical parameters: c, 58 × 1092, f39/1–2, × 50 bf. repository: gsc 138138. fig. 10. homotryblium abbreviatum, apical view. location: snorri j-90, 2249.45‒2266.6 m. sample: cs, p9747, 10. optical parameters: c, 50 × 1000, e29/2, × 50 bf. repository: gsc 138137. fig. 11. hystrichokolpoma cinctum, dorsal surface. location: karlsefni a-13, 2642.65–2651.79 m. sample: cs, p39613, 1. optical parameters: d, 128 × 920, m35/3, × 40 pc. repository: gsc 138041. fig. 12. hapsocysta? benteae. location: north leif i-05, 2310‒2320 m. sample: cs, yd17606, 4. optical parameters: e, 50.8 × 19.9, g50/3, × 60 bf. repository: mguh 31301. fig. 13. hystrichokolpoma globulus, left lateral view of left lateral surface. location: snorri j-90, 2293,65 m. sample: sw, p9370, 10. optical parameters: c, 181 × 1017, s31/0–4, × 50 bf. repository: gsc 138132. fig. 14. hystrichokolpoma globulus, left lateral view, same specimen as fig. 13, focussed on right lateral surface. location: snorri j-90, 2293,65 m. sample: sw, p9370, 10. optical parameters: c, 181 × 1017, s31/0–4, × 50 bf. repository: gsc 138132. fig. 15. hystrichokolpoma globulus, lateral view. location: snorri j-90, 2441.48‒2450.62 m. sample: cs, p9751, 10. optical parameters: c, 102 × 1010, k30/4, × 50 bf. repository: gsc 138139. fig. 16. heteraulacacysta porosa, antapical view of antapical surface. location: ralegh n-18, 1525 m. sample: cs, yd16219, 4. optical parameters: e, 36.3 × 20.5, g35/1, × 60 bf. repository: cnsopb. fig. 17. hystrichosphaeridium quadratum sp. nov., holotype, dorsal surface. location: gilbert f-53, 2770‒2780 m. sample: cs, p39505, 1. optical parameters: c, 113 × 992, l/29/3, × 50 bf. repository: gsc 137988. fig. 18. hystrichosphaeridium quadratum sp. nov., holotype, focussed on ventral surface. location: gilbert f-53, 2770‒2780 m. sample: cs, p39505, 1. optical parameters: c, 113 × 992, l29/3, × 50 bf. repository: gsc 137988. fig. 19. hystrichosphaeridium quadratum sp. nov., operculum, showing pre-apical process. location: karlsefni a-13, 3904.54–3913.68 m. sample: cs, p39658, 1. optical parameters: d, 125 × 999, m43/0‒3, × 40 pc. repository: gsc 138057. fig. 20. hystrichosphaeridium tubiferum, lateral view. location: gjoa o-37, 2120 m. sample: cs, yd16090, 4. optical parameters: e, 32.7 × 22.0, e31/0, × 60 bf. repository: cnsopb. plate 8 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 118 119 1 9 5 13 17 2 10 6 14 18 3 11 7 15 19 4 12 8 16 20 plate 8 plate 8 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 119 120 fig. 1. hystrichostrogylon digitus sp. nov., dorso-ventral view. location: karlsefni a-13, 2267.74–2276.88 m. sample: cs, p39599, 1. optical parameters: d, 200 × 878, u30/0, × 40 pc. repository: gsc 138030. fig. 2. hystrichostrogylon digitus sp. nov., dorso-ventral view. location: rut h-11, 2375‒2385 m. sample: cs, p39373, 1. optical parameters: d, 167 × 973, q40/3‒4, × 40 pc. repository: gsc 137950. fig. 3. hystrichostrogylon digitus sp. nov., holotype, dorso-ventral view. location: rut h-11, 2435‒2445 m. sample: cs, p39375, 1. optical parameters: d, 158 × 1050, q48/1, × 40 pc. repository: gsc 137952. fig. 4. hystrichosphaeropsis perforata, dorso-ventral view. location: skolp e-07, 1250 m. sample: cs, yd17614, 4. optical parameters: a, 43.3 × 101.4, p44/1, × 60 bf. repository: gsc 138164. fig. 5. impletosphaeridium apodastum sp. nov., holotype. location: hekja o-71, 4565‒4575 m. sample: cs, p18737, 1. optical parameters: c, 133 × 1080, n37/3, × 50 bf. repository: gsc 137903. fig. 6. impletosphaeridium apodastum sp. nov. location: karlsefni a-13, 3822.24–3831.38 m. sample: cs, p39655, 1. optical parameters: d, 201 × 990, u42/1‒3, × 40 pc. repository: gsc 138055. fig. 7. impletosphaeridium apodastum sp. nov. location: karlsefni a-13, 3989.88–3999.02 m. sample: cs, p39661, 1. optical parameters: d, 193 × 1070, t50/0, × 40 pc. repository: gsc 138063. fig. 8. impletosphaeridium apodastum sp. nov. location: karlsefni a-13, 3989.88–3999.02 m. sample: cs, p39661, 1. optical parameters: d, 173 × 899, r32/0, × 40 pc. repository: gsc 138062. fig. 9. impagidinium victorianum, ventral view of ventral surface. location: roberval k-92, 2270‒2280 m. sample: cs, p17700, 1. optical parameters: c, 66 × 1022, f32/3, × 50 bf. repository: gsc 137894. fig. 10. impagidinium victorianum, ventral view, same specimen as fig. 9, despite lower focus, ventral surface clear. location: roberval k-92, 2270‒2280 m. sample: cs, p17700, 1. optical parameters: c, 66 × 1022, f32/3, × 50 bf. repository: gsc 137894. fig. 11. impagidinium victorianum, ventral view, same specimen as fig. 9, focussed on periphery. location: roberval k-92, 2270‒2280 m. sample: cs, p17700, 1. optical parameters: c, 66 × 1022, f32/3, × 50 bf. repository: gsc 137894. fig. 12. hystrichosphaeropsis quasicribrata, dorso-ventral view. location: north leif i-05, 2550‒2560 m. sample: cs, yd17614, 4. optical parameters: e, 43.9 × 15.8, m43/2, × 60 bf. repository: gsc 138165. fig. 13. impagidinium victorianum, dorsal view of ventral surface. location: skolp e-07, 1175 m. sample: cs, yd15587, 3. optical parameters: a, 34.2 × 99.0, s435/1, × 60 bf. repository: mguh 31302. fig. 14. impagidinium victorianum, dorsal view of dorsal surface, same specimen as fig. 13. location: skolp e-07, 1175 m. sample: cs, yd15587, 3. optical parameters: a, 34.2 × 99.0, s435/1, × 60 bf. repository: mguh 31302. fig. 15. isabelidinium cretaceum, dorso-ventral view. location: north leif i-05, 2550‒2560 m. sample: cs, yd17614, 2. optical parameters: e, 17.4 × 20.3, g15/0, × 60 bf. repository: mguh 31303. fig. 16. isabelidinium cretaceum, dorso-ventral view. location: roberval k-92, 3060‒3070 m. sample: cs, p17726, 1. optical parameters: c, 157 × 1085, q38/2, × 50 bf. repository: gsc 137899. fig. 17. isabelidinium microarmum, dorsal view of dorsal surface. location: skolp e-07, 2285 m. sample: cs, yd15659, 2. optical parameters: a, 24.4 × 93.5, x24/2, × 60 bf. repository: mguh 31304. fig. 18. kleithriasphaeridium mantellii, right lateral view. location: maude bight, bylot island, section 3 (v002181), 1 m above base (sample pb-57). sample: os, p5189-46, h. optical parameters: c, 161 × 932, l54/1–3, × 50 bf. repository: gsc 138131. fig. 19. laciniadinium arcticum, ventral view of ventral surface. location: ogmund e-72, 1530 m (reworked specimen). sample: cs, yd15734, 3. optical parameters: a, 23.6 × 98.0, s23/4, × 60 bf. repository: mguh 31305. fig. 20. laciniadinium arcticum, dorsal view of ventral surface. location: skolp e-07, 1940 m. sample: cs, jeh15635, 3. optical parameters: a, 19.5 × 104.7, l19/3, × 60 bf. repository: mguh 31306. plate 9 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 120 121 1 9 5 13 17 2 10 6 14 18 3 11 7 15 19 4 12 8 16 20 plate 9 plate 9 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 121 122 fig. 1. lingulodinium funginum. location: karlsefni a-13, 2670.08–2679.22 m. sample: cs, p39614, 1. optical parameters: d, 145 × 1047, o48/3, × 40 pc. repository: gsc 138042. fig. 2. lingulodinium funginum. location: rut h-11, 3035‒3045 m. sample: cs, p39395, 1 optical parameters: d, 73 × 929, g36/1, × 40 pc. repository: gsc 137961. fig. 3. lingulodinium machaerophorum. location: rut h-11, 2585‒2595 m. sample: cs, p39380, 1. optical parameters: d, 105 × 900, k32/2, × 40 pc. repository: gsc 137954. fig. 4. lingulodinium machaerophorum, dorso-ventral view showing attached operculum. location: gjoa o-37, 1620 m. sample: cs, yd16074, 2. optical parameters: e, 18.0 × 19.6, h16/1, × 60 bf. repository: mguh 31307. fig. 5. nyktericysta davisii, dorso-ventral view. location: north leif i-05, 3180‒3190 m. sample: cs, yd17635, 4. optical parameters: e, 37.5 × 4.7, x36/0, × 40 bf. repository: gsc 138169. fig. 6. nyktericysta davisii, dorso-ventral view. location: north leif i-05, 3180‒3190 m. sample: cs, yd17635, 4. optical parameters: e, 20.9 × 12.4, p19/3, × 60 bf. repository: gsc 138168. fig. 7. nyktericysta dictyophora, dorso-ventral view. location: ogmund e-72, 1995 m. sample: cs, yd15765, 3. optical parameters: a, 37.3 × 94.1, w38/3, × 60 bf. repository: mguh 31308. fig. 8. nyktericysta dictyophora, dorso-ventral view. location: ogmund e-72, 1965 m. sample: cs, yd15763, 3. optical parameters: a, 23.8 × 100.8, p23/4, × 60 bf. repository: mguh 31309. fig. 9. lentinia serrata, dorsal view of dorsal surface. location: south labrador n-79, 2700‒2710 m. sample: cs, p39807, 1. optical parameters: c, 65 × 1027, f32/0–4, × 50 bf. repository: gsc 138122. fig. 10. palaeocystodinium golzowense, dorso-ventral view. location: rut h-11, 755‒765 m. sample: cs, p39319, 1. optical parameters: d, 61 × 1040, e47/3, × 40 pc. repository: gsc 137920. fig. 11. nyktericysta tripenta, dorso-ventral view; endophragm visible apically. location: ogmund e-72, 2550 m. sample: cs, jeh15891, 3. optical parameters: a, 47.9 × 105.8, k49/3, × 60 bf. repository: mguh 31310. fig. 12. nyktericysta tripenta, dorso-ventral view, same specimen as fig. 11. location: ogmund e-72, 2550 m. sample: cs, jeh15891, 3. optical parameters: a, 47.9 × 105.8, k49/3, × 60 bf. repository: mguh 31310. fig. 13. odontochitina porifera, dorsal view. location: skolp e-07, 1895 m. sample: cs, jeh15633, 4. optical parameters: a, 33.2 × 107.7, h33/4, × 60 bf. repository: mguh 31311. fig. 14. odontochitina ancala, dorsal view. location: north leif i-05, 2790‒2800 m. sample: cs, yd17622, 3. optical parameters: e, 20.1 × 18.5, j18/1, × 40 bf. repository: mguh 31312. fig. 15. odontochitina costata, dorsal view. location: skolp e-07, 2345 m. sample: cs, yd15663, 2. optical parameters: a, 40.3 × 111.8, d40/2, × 60 bf. repository: mguh 31313. fig. 16. licracysta corymbus, dorsal surface. location: rut h-11, 1205‒1215 m. sample: cs, p39334, 1. optical parameters: d, 150 × 974, p40/1‒2, × 40 pc. repository: gsc 137931. fig. 17. oligosphaeridium albertense. location: rut h-11, 1775‒1785 m. sample: cs, p39353, 1. optical parameters: d, 85 × 866, h29/0, × 40 pc. repository: gsc 137939. fig. 18. oligosphaeridium pulcherrimum, dorso-ventral view. location: gilbert f-53, 2705‒2715 m. sample: cs, p39503, 1. optical parameters: c, 24 × 1012, y31/3, × 50 bf. repository: gsc 137986. fig. 19. oligosphaeridium pulcherrimum. location: rut h-11, 1745‒1755 m. sample: cs, p39352, 1. optical parameters: d, 168 × 880, q30/4, × 40 pc. repository: gsc 137938. fig. 20. oligosphaeridium totum. location: rut h-11, 2225‒2235 m. sample: cs, p39368, 1. optical parameters: d, 154 × 1052, p48/0, × 40 pc. repository: gsc 137944. plate 10 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 122 123 1 9 5 13 17 2 10 6 14 18 3 11 7 15 19 4 12 8 16 20 plate 10 plate 10 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 123 124 fig. 1. palaeocystodinium bulliforme, right lateral view. location: gilbert f-53, 2470‒2480 m. sample: cs, p19228, 1. optical parameters: c, 163 × 1082, q28/3, × 50 bf. repository: gsc 137904. fig. 2. palaeocystodinium bulliforme. location: ogmund e-72, 1545 m. sample: cs, yd15735, 4. optical parameters: a, 30.6 × 110.6, e31/3, × 40 bf. repository: cnlopb. fig. 3. palaeocystodinium teespinosum, left lateral view. location: karlsefni a-13, 969.28–978.42 m. sample: cs, p39553, 1. optical parameters: d, 201 × 922, u35/1, × 40 pc. repository: gsc 138007. fig. 4. palaeohystrichophora infusorioides, dorsal view. location: north leif i-05, 3000‒3010 m. sample: cs, yd17629, 3. optical parameters: e, 25.0 × 20.2, g23/0, × 60 bf. repository: mguh 31314. fig. 5. operculodinium centrocarpum. location: karlsefni a-13, 694.95‒704.10 m. sample: cs, p39543, 1. optical parameters: d, 143 × 996, o42/0, × 40 pc. repository: gsc 137996. fig. 6. operculodinium centrocarpum, operculum. location: karlsefni a-13, 722.38‒731.53 m. sample: cs, p39544, 1. optical parameters: d, 169 × 934, r36/1‒2, × 40 pc. repository: gsc 137998. fig. 7. palaeoperidinium pyrophorum, dorsal surface. location: gilbert f-53, 2680‒2690 m. sample: cs, p39502, 1. optical parameters: c, 68 × 1101, f40/3‒g40/1, × 50 bf. repository: gsc 137985. fig. 8. palaeoperidinium pyrophorum, ventral view. location: roberval k-92, 3070‒3080 m. sample: cs, p2008175, 1. optical parameters: c, 200 × 982, u27/4‒u28/3, × 50 bf. repository: gsc 137907. fig. 9. palynodinium grallator, apical view. location: roberval k-92, 2840‒2850 m. sample: cs, p17719, 1. optical parameters: c, 30 × 1087, c38/2–c39/1, × 50 bf. repository: gsc 137898. fig. 10. palynodinium grallator, dorsal view. location: north leif i-05, 2520‒2530 m. sample: cs, yd17613, 4. optical parameters: e, 44.0 × 16.5, l46/3, × 60 bf. repository: gsc 138163. fig. 11. petalodinium condylos, ventral view of dorsal surface. location: roberval k-92, 2420‒2430 m. sample: cs, p17705, 1. optical parameters: c, 110 × 1018, q31/2–4, × 50 bf. repository: gsc 137896. fig. 12. petalodinium condylos, ventral view of dorsal surface. location: roberval k-92, 2240‒2250 m. sample: cs, p17699, 1. optical parameters: c, 175 × 1052, s35/1, × 50 bf. repository: gsc 137892. fig. 13. phthanoperidinium coreoides, dorsal surface. location: rut h-11, 1145‒1155 m. sample: cs, p39332, 1. optical parameters: d, 126 × 991, m42/0, × 40 pc. repository: gsc 137928. fig. 14. phthanoperidinium coreoides, dorsal surface. location: north leif i-05, 720‒730 m. sample: cs, yd17553, 2. optical parameters: e, 46.7 × 11.2, q46/1, × 60 bf. repository: mguh 31315. fig. 15. phthanoperidinium levimurum. location: karlsefni a-13, 3273.59–3282.74 m. sample: cs, p39636, 1. optical parameters: d, 100 × 1064, j50/3, × 40 pc. repository: gsc 138050. fig. 16. phthanoperidinium multispinum, dorsal surface. location: south labrador n-79, 1260‒1270 m. sample: cs, p39762, 1. optical parameters: c, 87 × 1017, j31/2, × 50 bf. repository: gsc 138110. fig. 17. phthanoperidinium multispinum, ventral surface. location: karlsefni a-13, 1956.84–1965.98 m. sample: cs, p39588, 1. optical parameters: d, 194 × 1025, t45/4, × 40 pc. repository: gsc 138020. fig. 18. phthanoperidinium regale, ventral view of ventral surface. location: hellefisk-1, 1289 m. sample: cs, 02b2097-4, 4. optical parameters: a, 48.5 × 105.9, k49/4, × 40 bf. repository: mguh 26504. fig. 19. phthanoperidinium regale, ventral view of dorsal surface, same specimen as fig. 18. location: hellefisk-1, 1289 m. sample: cs, 02b2097-4, 4. optical parameters: a, 48.5 × 105.9, k49/4, × 60 bf. repository: mguh 26504. fig. 20. phthanoperidinium stockmansii, ventral view of ventral surface. location: south labrador n-79, 1230‒1240 m. sample: cs, p39761, 1. optical parameters: c, 67 × 1021, f32/3, × 50 bf. repository: gsc 138109. plate 11 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 124 125 1 9 5 13 17 2 10 6 14 18 3 11 7 15 19 4 12 8 16 20 plate 11 plate 11 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 125 126 fig. 1. phelodinium kozlowskii, dorsal surface. location: gilbert f-53, 2740‒2750 m. sample: cs, p39504, 1. optical parameters: c, 203 × 1020, u31/4–u32/3, × 50 bf. repository: gsc 137987. fig. 2. phelodinium kozlowskii, oblique left lateral view. location: bjarni o-82, 2175‒2185 m. sample: cs, p39727, 1. optical parameters: c, 139 × 993, o29/1, × 50 bf. repository: gsc 138072. fig. 3. phelodinium kozlowskii, dorsal view showing dorsal and ventral surfaces. location: south labrador n-79, 3030‒3040 m. sample: cs, p39818, 1. optical parameters: c, 218 × 1000, n29/2, × 50 bf. repository: gsc 138123. fig. 4. pseudoceratium sp., ventral view. location: roberval k-92, 3070‒3080 m. sample: cs, p2008175, 1. optical parameters: c, 210 × 993, v29/0‒3, × 40 pc. repository: gsc 137908. fig. 5. piladinium columna, ventral view of dorsal surface. location: bjarni o-82, 1785‒1795 m. sample: cs, p39714, 1. optical parameters: c, 221 × 1054, w32/3–4, × 50 bf. repository: gsc 138069. fig. 6. piladinium columna, dorsal view of dorsal surface. location: roberval k-92, 2270‒2280 m. sample: cs, p17700, 1. optical parameters: c, 200 × 1014, u31/0, × 50 bf. repository: gsc 137893. fig. 7. piladinium edwardsii, dorsal view of dorsal surface. location: ogmund e-72, 1305 m. sample: sw, jeh15551, 4. optical parameters: a, 27.3 × 98.6, t27/1, × 60 bf. repository: mguh 31316. fig. 8. piladinium edwardsii, dorsal view of dorsal surface. location: north leif i-05, 2010‒2020 m. sample: cs, yd17596, 3. optical parameters: e, 21.0 × 10.6, r18/1, × 60 bf. repository: mguh 31317. fig. 9. raphidodinium fucatum, dorsal surface. location: bjarni o-82, 2415‒2425 m. sample: cs, p39735, 1. optical parameters: c, 228 × 1077, x32/0–2, × 50 bf. repository: gsc 138074. fig. 10. raphidodinium fucatum. location: skolp e-07, 1475 m. sample: cs, yd15606, 3. optical parameters: a, 43.7 × 94.7, w44/1, × 60 bf. repository: mguh 31318. fig. 11. reticulatosphaera actinocoronata. location: karlsefni a-13, 722.38‒731.53 m. sample: cs, p39544, 1. optical parameters: d, 146 × 930, o36/3, × 40 pc. repository: gsc 137997. fig. 12. reticulatosphaera actinocoronata. location: rut h-11, 755‒765 m. sample: cs, p39319, 1. optical parameters: d, 89 × 1006, h43/4, × 40 pc. repository: gsc 137921. fig. 13. rhombodinium draco, ventral view of dorsal surface. location: snorri j-90, 1673.37‒1682.52 m. sample: cs, p9729, 10. optical parameters: c, 93 × 1007, j29/0, × 40 pc. repository: gsc 138134. fig. 14. rhombodinium draco, ventral view of ventral surface. location: bjarni o-82, 615‒625 m. sample: cs, p39675, 1. optical parameters: c, 148 × 1035, p33/0, × 40 bf. repository: gsc 138067. fig. 15. scalenodinium scalenum sp. nov., holotype, left lateral view. location: gilbert f-53, 2135‒2145 m. sample: cs, p38484, 1. optical parameters: c, 193 × 1065, t37/4, × 50 bf. repository: gsc 137916. fig. 16. scalenodinium scalenum sp. nov., left lateral view. location: gilbert f-53, 2135‒2145 m. sample: cs, p39484, 1. optical parameters: c, 48 × 1020, d32/3–d31/4, × 50 bf. repository: gsc 137982. fig. 17. scalenodinium scalenum sp. nov. location: ogmund e-72, 1391 m. sample: sw, yd15553, 2. optical parameters: a, 26.8 × 107.0, j26/4, × 60 bf. repository: mguh 31319. fig. 18. scalenodinium scalenum sp. nov. location: ralegh n-18, 3005 m. sample: cs, yd16263, 4. optical parameters: e, 41.6 × 11.1, q40/4, × 60 bf. repository: mguh 31320. fig. 19. scalenodinium scalenum sp. nov., oblique left lateral view. location: ralegh n-18, 3045 m. sample: cs, yd16264, 3. optical parameters: e, 28.9 × 19.9, g27/0, × 60 bf. repository: mguh 31321. fig. 20. scalenodinium scalenum sp. nov. location: north leif i-05, 2130‒2140 m. sample: cs, yd17600, 3. optical parameters: e, 29.4 × 7.6, u27/2, × 60 bf. repository: gsc 138156. plate 12 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 126 127 1 9 5 13 17 2 10 6 14 18 3 11 7 15 19 4 12 8 16 20 plate 12 plate 12 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 127 128 fig. 1. rottnestia borussica, right lateral view. location: south labrador n-79, 2430‒2440 m. sample: cs, p39799, 1. optical parameters: c, 105 × 990, k28/4, × 50 bf. repository: gsc 138121. fig. 2. schematophora speciosa, dorsal surface. location: south labrador n-79, 1290‒1300 m. sample: cs, p39763, 1. optical parameters: c, 126 × 982, n28/1–n27/2, × 50 bf. repository: gsc 138111. fig. 3. senegalinium iterlaaense, dorsal surface. location: gjoa o-37, 3560 m. sample: cs, yd16131, 5. optical parameters: e, 46.4 × 11.6, q45/4, × 60 bf. repository: gsc 138146. fig. 4. senegalinium iterlaaense, ventral view of dorsal surface. location: gjoa o-37, 3610 m. sample: cs, yd16132, 3. optical parameters: e, 37.1 × 3.9, x36/3, × 60 bf. repository: mguh 31322. fig. 5. senoniasphaera inornata, dorsal view. location: north leif i-05, 2490‒2500 m. sample: cs, yd17612, 3. optical parameters: e, 41.2 × 8.3, t40/2, × 60 bf. repository: gsc 138162. fig. 6. senoniasphaera microreticulata, dorsal view. location: north leif i-05, 2820‒2830 m. sample: cs, yd17623, 2. optical parameters: e, 41.5 × 9.2, s40/2, × 60 bf. repository: mguh 31323. fig. 7. senoniasphaera rotundata, dorsal view of dorsal surface. location: south labrador n-79, 3480‒3490 m. sample: cs, p39833, 1. optical parameters: c, 114 × 1037, l33/4, × 50 bf. repository: gsc 138126. fig. 8. sophismatia tenuivirgula, dorso-ventral view. location: north leif i-05, 1560‒1570 m. sample: cs, yd17581, 2. optical parameters: e, 49.0 × 14.2, n48/2, × 60 bf. repository: mguh 31324. fig. 9. simplicidinium insolitum. location: rut h-11, 1205‒1215 m. sample: cs, p39334, 1. optical parameters: d, 152 × 871, p29/2, × 40 pc. repository: gsc 137932. fig. 10. simplicidinium insolitum. location: karlsefni a-13, 804.68–813.83 m. sample: cs, p39547, 1. optical parameters: d, 162 × 1052, q48/0, × 40 pc. repository: gsc 137999. fig. 11. spinidinium echinoideum, dorsal view of ventral surface. location: ogmund e-72, 1605 m. sample: cs, yd15739, 3. optical parameters: a, 37.3 × 108.0, h38/1, × 60 bf. repository: mguh 31325. fig. 12. spinidinium echinoideum, dorsal view, same specimen as fig. 11, focussed on ambitus. location: ogmund e-72, 1605 m. sample: cs, yd15739, 3. optical parameters: a, 37.3 × 108.0, h38/1, × 60 bf. repository: mguh 31325. fig. 13. spiniferites ovatus, right lateral view. location: rut h-11, 755‒765 m. sample: cs, p39319, 1. optical parameters: d, 135 × 933, n36/0, × 40 pc. repository: gsc 137923. fig. 14. spiniferites pseudofurcatus. location: rut h-11, 2285‒2295 m. sample: cs, p39370, 1. optical parameters: d, 128 × 956, m38/4, × 40 pc. repository: gsc 137947. fig. 15. spiniferites scabrosus. location: roberval k-92, 3010 m. sample: sw, p2008173, 1. optical parameters: c, 210 × 1088, v38/3–4, × 50 bf. repository: gsc 137906. fig. 16. spongodinium delitiense, apical view. location: skolp e-07, 1070 m. sample: cs, yd15580, 4. optical parameters: a, 46.1× 96.5, u47/1, × 60 bf. repository: gsc 138143. fig. 17. stichodinium lineidentatum, dorsal view of ventral surface. location: nukik-2, 1862 m. sample: cs, 05b2163, 3. optical parameters: a, 43.6 × 111.4, d44/4, × 60 bf. repository: mguh 31326. fig. 18. stichodinium lineidentatum, dorsal view of dorsal surface, same specimen as fig. 17. location: nukik-2, 1862 m. sample: cs, 05b2163, 3. optical parameters: a, 43.6 × 111.4, d44/4, × 60 bf. repository: mguh 31326. fig. 19. spongodinium obscurum, dorsal surface. location: ggu 400712 borehole, 62.13‒62.25 m. sample: cc, d309-f, 3. optical parameters: a, 32.0 × 109.0, g32/3, × 60 bf. repository: mguh 31327. fig. 20. spongodinium delitiense, right lateral view. location: north leif i-05, 2460‒2470 m. sample: cs, yd17611, 3. optical parameters: e, 29.2 × 18.0, j27/2, × 60 bf. repository: mguh 31328. plate 13 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 128 129 32 4 17 1 9 5 13 10 6 14 18 11 7 15 19 12 8 16 20 plate 13 plate 13 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 129 130 fig. 1. spongodinium grossum, oblique left lateral view. location: skolp e-07, 1565 m. sample: cs, jeh15612, 3. optical parameters: a, 29.5 × 110.6, e29/4, × 60 bf. repository: mguh 31329. fig. 2. spongodinium grossum, left lateral view. location: skolp e-07, 1565 m. sample: cs, jeh15612, 5. optical parameters: a, 27.3 × 98.5, s27/2, × 60 bf. repository: mguh 31330. fig. 3. spongodinium grossum, ventral surface. location: bjarni o-82, 2445‒2455 m. sample: cs, p39736, 1. optical parameters: c, 181 × 991, s29/3, × 40 bf. repository: gsc 138075. fig. 4. spongodinium grossum, same specimen as fig. 3, focussed on dorsal surface. location: bjarni o-82, 2445‒2455 m. sample: cs, p39736, 1. optical parameters: c, 181 × 991, s29/3, × 40 bf. repository: gsc 138075. fig. 5. subtilisphaera perlucida, dorsal surface. location: north leif i-05, 2910‒2920 m. sample: cs, yd17626, 4. optical parameters: e, 23.4 × 19.2, h21/0, × 60 bf. repository: gsc 138167. fig. 6. surculosphaeridium convocatum sp. nov., ventral surface. location: south labrador n-79, 3510‒3520 m. sample: cs, p39834, 1. optical parameters: c, 148 × 982, p28/1, × 50 bf. repository: gsc 138127. fig. 7. surculosphaeridium convocatum sp. nov., holotype, dorsal surface. location: south labrador n-79, 3510‒3520 m. sample: cs, p39834, 1. optical parameters: c, 78 × 991, g29/3–h28/2, × 50 bf. repository: gsc 138128. fig. 8. surculosphaeridium convocatum sp. nov., lateral view. location: north leif i-05, 2760‒2775 m. sample: cs, yd17611, 2. optical parameters: e, 21.9 × 19.4, h20/1, × 60 bf. repository: mguh 31331. fig. 9. talladinium? clathratum, dorsal view, focussed on ambitus. location: ralegh n-18, 1445 m. sample: cs, yd16217, 2. optical parameters: e, 37.6 × 14.5, m36/4, × 40 bf. repository: mguh 31332. fig. 10. talladinium? clathratum, dorsal view, same specimen as fig. 9. location: ralegh n-18, 1445 m. sample: cs, yd16217, 2. optical parameters: e, 37.6 × 14.5, m36/4, × 60 bf. repository: mguh 31332. fig. 11. talladinium pellis sp. nov., holotype, dorsal view of dorsal surface. location: gjoa o-37, 1840 m. sample: cs, yd16082, 3. optical parameters: c, 21.9 × 22.0, e20/3, × 60 bf. repository: gsc 138082. fig. 12. talladinium pellis sp. nov., holotype, dorsal view of ventral surface, same specimen as fig. 11. location: gjoa o-37, 1840 m. sample: cs, yd16082, 3. optical parameters: c, 21.9 × 22.0, e20/3, × 60 bf. repository: gsc 138082. fig. 13. taurodinium granulatum sp. nov., holotype, dorso-ventral view. location: ikermiut-1, 2340 m. sample: cs, 2289 s-261, 3. optical parameters: b, 17.82 × 9.01, t19/1, × 60 bf. repository: mguh 31333. fig. 14. taurodinium granulatum sp. nov., dorso-ventral view. location: ralegh n-18, 3045 m. sample: cs, yd16264, 3. optical parameters: e, 23.6 × 14.8, m21/4, × 60 bf. repository: mguh 31334. fig. 15. taurodinium granulatum sp. nov., dorso-ventral view. location: ralegh n-18, 3405 m. sample: cs, yd16341, 3. optical parameters: e, 20.8 × 16.9, k19/3, × 60 bf. repository: mguh 31335. fig. 16. talladinium pellis sp. nov., ventral view. location: gjoa o-37, 1840 m. sample: cs, yd16082, 4. optical parameters: c, 18.56 × 15.03, l20/1, × 60 bf. repository: mguh 31336. fig. 17. taurodinium granulatum sp. nov., dorso-ventral view. location: ralegh n-18, 3425 m. sample: cs, yd16342, 4. optical parameters: e, 33.0 × 18.6, j31/2, × 60 bf. repository: gsc 138147. fig. 18. taurodinium granulatum sp. nov. location: hekja o-71, 3120 m. sample: cs, jeh16030, 2. optical parameters: e, 24.9 × 23.6, c23/3, × 60 bf. repository: mguh 31337. fig. 19. tenua hystrix, ventral view of dorsal surface. location: roberval k-92, 3090‒3100 m. sample: cs, p17727, 1. optical parameters: c, 175 × 1086, s38/2, × 50 bf. repository: gsc 137900. fig. 20. tenua hystrix, lateral view. location: two snout creek, bylot island, section 538080e, 105 m above base (sample hfb-09-42). sample: os, p514833, d. optical parameters: c, 145 × 918, m55/3–4, × 50 bf. repository: gsc 138129. plate 14 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 130 131 1 9 5 13 17 2 10 6 14 18 3 11 7 15 19 4 12 8 16 20 plate 14 plate 14 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 131 132 fig. 1. thalassiphora delicata, dorsal surface. location: north leif i-05, 2370‒2380 m. sample: cs, yd17608, 4. optical parameters: e, 35.7 × 12.9, p34/1, × 60 bf. repository: mguh 31338. fig. 2. thalassiphora fenestrata, right lateral view. location: ralegh n-18, 1445 m. sample: cs, yd16217, 2. optical parameters: e, 32.5 × 20.5, f31/4, × 60 bf. repository: mguh 31339. fig. 3. thalassiphora fenestrata, left lateral view. location: ralegh n-18, 1445 m. sample: cs, yd16217, 2. optical parameters: e, 41.1 × 11.2, q40/2, × 60 bf. repository: mguh 31340. fig. 4. thalassiphora fenestrata, dorso-ventral view. location: ralegh n-18, 1605 m. sample: cs, yd16221, 4. optical parameters: e, 32.8 × 6.2, v31/3, × 40 bf. repository: cnsopb. fig. 5. thalassiphora pelagica, dorso-ventral view. location: rut h-11, 2765‒2775 m. sample: cs, p39386, 1. optical parameters: d, 181 × 906, s33/0, × 40 pc. repository: gsc 137956. fig. 6. thalassiphora pelagica, dorsal surface. location: karlsefni a-13, 2615.22–2624.36 m. sample: cs, p39612, 1. optical parameters: d, 52 × 902, d33/3, × 40 pc. repository: gsc 138039. fig. 7. tanyosphaeridium xanthiopyxides. location: gilbert f-53, 2590‒2600 m. sample: cs, p39499, 1. optical parameters: c, 94 × 1004, j30/0, × 50 bf. repository: gsc 137984. fig. 8. trichodinium castanea, oblique left lateral view. location: skolp e-07, 1430 m. sample: cs, yd15603, 3. optical parameters: a, 36.0 × 100.7, q356/2, × 60 bf. repository: mguh 31341. fig. 9. trithyrodinium? conservatum sp. nov., dorso-ventral view, with focus on ambitus. location: ikermiut-1, 1155 m. sample: sw, 04e006504, 3. optical parameters: a, 33.0 × 112.8, c33/3, × 60 bf. repository: mguh 26500. fig. 10. trithyrodinium? conservatum sp. nov., dorso-ventral view. location: ikermiut-1, 1155 m. sample: sw, 04e006504, 2. optical parameters: a, 41.0 × 113.5, c41/2, × 60 bf. repository: mguh 26501. fig. 11. trithyrodinium? conservatum sp. nov., dorso-ventral view, same specimen as fig. 10, focussed on dorsal surface. location: ikermiut-1, 1155 m. sample: sw, 04e006504, 2. optical parameters: a, 41.0 × 113.5, c41/2, × 60 bf. repository: mguh 26501. fig. 12. trithyrodinium? conservatum sp. nov., holotype, dorsal surface. location: ikermiut-1, 1155 m. sample: sw, 04e006504, 2. optical parameters: a, 36.0 × 97.1, u36/1, × 60 bf. repository: mguh 26502. fig. 13. trithyrodinium? conservatum sp. nov., dorsal surface. location: ralegh n-18, 1485 m. sample: cs, yd16218, 2. optical parameters: e, 30.3 × 23.1, d29/1, × 60 bf. repository: mguh 31342. fig. 14. trithyrodinium? conservatum sp. nov., dorsal surface. location: ralegh n-18, 1525 m. sample: cs, yd16219, 2. optical parameters: e, 28.7 × 5.8, w27/2, × 60 bf. repository: mguh 31343. fig. 15. trithyrodinium suspectum, dorso-ventral view. location: skolp e-07, 2375 m. sample: cs, yd15665, 4. optical parameters: a, 39.0 × 107.4, h39/4, × 60 bf. repository: mguh 31344. fig. 16. trithyrodinium quinqueangulare, oblique dorso ventral view. location: south labrador n-79, 3300‒3310 m. sample: cs, p39827, 1. optical parameters: c, 223 × 1020, w32–x31, × 50 bf. repository: gsc 138125. fig. 17. trithyrodinium evittii, oblique apical view. location: north leif i-05, 2370–2380 m. sample: cs, yd17608, 3. optical parameters: e, 24.5 × 11.0, q23/3, × 60 bf. repository: mguh 31345. fig. 18. trithyrodinium evittii, dorsal surface. location: gilbert f-53, 2920‒2930 m. sample: cs, p39510, 1. optical parameters: c, 100 × 1067, k36/3–4, × 50 bf. repository: gsc 137989. fig. 19. trithyrodinium evittii, apical view. location: roberval k-92, 3090‒3100 m. sample: cs, p17727, 1. optical parameters: c, 197 × 1095, u39/0‒2, × 50 bf. repository: gsc 137901. fig. 20. tuberculodinium vancampoae, apical view. location: karlsefni a-13, 557.79–566.93 m. sample: cs, p39538, 1. optical parameters: d, 146 × 1038, p46/2, × 40 pc. repository: gsc 137993. plate 15 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 132 133 1 9 5 13 17 2 10 6 14 18 3 11 7 15 19 4 12 8 16 20 plate 15 plate 15 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 133 134 fig. 1. vesperopsis longicornis, dorso-ventral view. location: north leif i-05, 3360–3370 m. sample: cs, yd17641, 4. optical parameters: e, 55.6 × 10.1, r55/1, × 60 bf. repository: mguh 31346. fig. 2. vesperopsis longicornis, dorso-ventral view. location: north leif i-05, 3360–3370 m. sample: cs, yd17641, 3. optical parameters: e, 35.2 × 23.4, c34/1, × 60 bf. repository: mguh 31347. fig. 3. wetzeliella articulata, dorsal view. location: rut h-11, 3725‒3735 m. sample: cs, p39417, 1. optical parameters: d, 152 × 847, p27/0, × 40 pc. repository: gsc 137971. fig. 4. wetzeliella articulata, dorsal view. location: rut h-11, 3095‒3105 m. sample: cs, p39397, 1. optical parameters: d, 81 × 949, g38/3, × 40 pc. repository: gsc 137963. fig. 5. wallodinium luna, lateral view. location: skolp e-07, 1610 m. sample: cs, jeh15615, 3. optical parameters: a, 39.6 × 100.2, r40/1, × 60 bf. repository: mguh 31348. fig. 6. xenascus wetzelii, dorsal view of dorsal surface. location: skolp e-07, 1475 m. sample: cs, yd15606, 3. optical parameters: a, 41.0 × 111.2, e41/2, × 40 bf. repository: mguh 31349. fig. 7. xenascus wetzelii, dorsal view, same specimen as fig. 6, showing enlarged view of ventral surface. location: skolp e-07, 1475 m. sample: cs, yd15606, 3. optical parameters: a, 41.0 × 111.2, e41/2, × 60 bf. repository: mguh 31349. fig. 8. xenascus wetzelii, dorsal view, same specimen as fig. 6, showing enlarged view of dorsal surface. location: skolp e-07, 1475 m. sample: cs, yd15606, 3. optical parameters: a, 41.0 × 111.2, e41/2, × 60 bf. repository: mguh 31349. fig. 9. xenascus ceratioides, ventral view. location: rut h-11, 2255‒2235 m. sample: cs, p39369, 1. optical parameters: d, 116 × 954, l38/0, × 40 pc. repository: gsc 137945. fig. 10. xenascus ceratioides, dorsal view. location: skolp e-07, 2435 m. sample: cs, yd15669, 3. optical parameters: a, 24.6 × 107.9, h24/2, × 60 bf. repository: gsc 138144. fig. 11. fromea quadrangularis sp. nov. location: skolp e-07, 2075 m. sample: cs, jeh15643, 3. optical parameters: a, 42.1 × 94.3, w43/3, × 60 bf. repository: mguh 31350. fig. 12. fromea quadrangularis sp. nov. location: skolp e-07, 2090 m. sample: cs, jeh15644, 5. optical parameters: a, 44.2 × 109.3, g45/1, × 60 bf. repository: gsc 137882. fig. 13. microsphaeridium ancistroides. location: rut h-11, 665‒675 m. sample: cs, p39316, 1. optical parameters: d, 108 × 900, k33/3, × 40 pc. repository: gsc 137917. fig. 14. microsphaeridium ancistroides. location: rut h-11, 1835‒1845 m. sample: cs, p39355, 1. optical parameters: d, 186 × 907, s33/3‒4, × 40 pc. repository: gsc 137941. fig. 15. fromea quadrangularis sp. nov. location: skolp e-07, 2360 m. sample: cs, yd15664, 2. optical parameters: a, 47.6 × 99.1, r45/4, × 60 bf. repository: mguh 31351. fig. 16. fromea quadrangularis sp. nov., holotype. location: skolp e-07, 2375 m. sample: cs, yd15665, 4. optical parameters: a, 30.9 × 111.9, d31/1, × 60 bf. repository: mguh 31352. fig. 17. microsphaeridium ancistroides. location: karlsefni a-13, 667.52–676.66 m. sample: cs, p39542, 1. optical parameters: d, 170 × 946, r37/2, × 40 pc. repository: gsc 137994. fig. 18. microsphaeridium ancistroides. location: rut h-11, 665‒675 m. sample: cs, p39316, 1. optical parameters: d, 135 × 1003, n43/0, × 40 pc. repository: gsc 137918. fig. 19. fromea nicosia. location: skolp e-07, 2360 m. sample: cs, yd15664, 2. optical parameters: a, 25.5 × 93.1, x25/0, × 60 bf. repository: mguh 31353. fig. 20. palambages sp. location: rut h-11, 2405‒2415 m. sample: cs, p39374, 1. optical parameters: d, 169 × 1003, r43/1‒2, × 40 pc. repository: gsc 137951. plate 16 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 134 135 1 9 5 13 17 2 10 6 14 18 3 11 7 15 19 4 12 8 16 20 plate 16 plate 16 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 135 136 fig. 1. paralecaniella indentata. location: karlsefni a-13, 2176.30–2185.44 m. sample: cs, p39596, 1. optical parameters: d, 162 × 873, q30/1, × 40 pc. repository: gsc 138024. fig. 2. paralecaniella indentata. location: hellefisk-1, 1588 m. sample: sw, 02e6529-3, 3. optical parameters: a, 30.5 × 106.3, k30/0, × 60 bf. repository: mguh 26544. fig. 3. pediastrum sp. location: rut h-11, 1865‒1875 m. sample: cs, p39356, 1. optical parameters: d, 136 × 1057, n49/3, × 40 pc. repository: gsc 137943. fig. 4. pediastrum sp. location: karlsefni a-13, 2286.03–2295.17 m. sample: cs, p39600, 1. optical parameters: d, 167 × 1007, q44/3, × 40 pc. repository: gsc 138032. fig. 5. tetraporina sp. a. location: gjoa o-37, 1680 m. sample: cs, yd16076, 2. optical parameters: e, 23.1 × 14.4, n21/0, × 60 bf. repository: mguh 31354. fig. 6. tetraporina sp. a. location: north leif i-05, 540‒550 m. sample: cs, yd17547, 3. optical parameters: e, 37.0 × 11.3, q35/4, × 60 bf. repository: gsc 138148. fig. 7. tetraporina sp. a. location: karlsefni a-13, 2011.70–2020.85 m. sample: cs, p39590, 1. optical parameters: d, 191 × 923, t35/0, × 40 pc. repository: gsc 138021. fig. 8. tetraporina sp. b. location: rut h-11, 3035‒3045 m. sample: cs, p39395, 1. optical parameters: d, 134 × 931, n36/0, × 40 pc. repository: gsc 137962. fig. 9. tetraporina sp. b. location: gjoa o-37, 1620 m. sample: cs, yd16074, 4. optical parameters: a, 42.5 × 98.2, s43/0, × 60 bf. repository: mguh 31355. fig. 10. tetraporina sp. b. location: gjoa o-37, 1620 m. sample: cs, yd16074, 4. optical parameters: a, 34.0 × 105.0, l34/0, × 60 bf. repository: mguh 31356. fig. 11. tetraporina sp. b. location: skolp e-07, 935 m. sample: cs, yd15568, 4. optical parameters: a, 22.5 × 98.4, s22/2, × 60 bf. repository: gsc 138141. fig. 12. tetraporina sp. b. location: rut h-11, 3485‒3495 m. sample: cs, p39410, 1. optical parameters: d, 164 × 904, q33/1‒3, × 40 pc. repository: gsc 137968. fig. 13. translucentipollis contiguus. location: rut h-11, 3785‒3795 m. sample: cs, p39419, 1. optical parameters: d, 155 × 978, p40/4, × 40 pc. repository: gsc 137973. fig. 14. aquilapollenites quadrilobus. location: karlsefni a-13, 1956.84–1965.98 m. sample: cs, p39588, 1. optical parameters: d, 138 × 1037, n47/3, × 40 pc. repository: gsc 138019. fig. 15. aquilapollenites quadrilobus. location: bjarni o-82, 799‒800 m. sample: cs, p39681, 1. optical parameters: c, 220 × 1048, w34/4, × 40 bf. repository: gsc 138068. fig. 16. parviprojectus reticulatus. location: rut h-11, 3605‒3615 m. sample: cs, p39413, 1. optical parameters: d, 198 × 855, u28/1, × 40 pc. repository: gsc 137969. fig. 17. aquilapollenites quadrilobus. location: hekja o-71, 2080 m. sample: cs, jeh16005, 1. optical parameters: c, 145 × 1007, o30/4–p30/2, × 50 bf. repository: gsc 137884. fig. 18. afropollis sp. location: ogmund e-72, 1845 m. sample: cs, yd15755, 3. optical parameters: a, 26.4 × 101.8, o26/4, × 60 bf. repository: mguh 31357. fig. 19. afropollis sp. location: ogmund e-72, 1845 m. sample: cs, yd15755, 3. optical parameters: a, 26.8 × 96.7, u27/1, × 60 bf. repository: mguh 31358. fig. 20. appendicisporites potomacensis. location: ogmund e-72, 2340 m. sample: cs, jeh15877, 3. optical parameters: a, 16.3 × 111.3, d16/3, × 60 bf. repository: mguh 31359. plate 17 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 136 137 1 9 5 13 17 2 10 6 14 18 3 11 7 15 19 4 12 8 16 20 plate 17 plate 17 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 137 138 fig. 1. appendicisporites unicus. location: ogmund e-72, 2205 m. sample: cs, jeh15868, 3. optical parameters: a, 28.6 × 102.6, n28/4, × 60 bf. repository: gsc 137883. fig. 2. azolla sp. location: ralegh n-18, 2885 m. sample: cs, yd16260, 3. optical parameters: e, 25.1 × 8.1, t23/2, × 60 bf. repository: mguh 31360. fig. 3. azolla sp. location: gjoa o-37, 1858‒1868 m. sample: cs, p16963, 1. optical parameters: c, 50 × 1011, e31/1, × 50 bf. repository: gsc 137886. fig. 4. azolla sp. location: karlsefni a-13, 2395.76–2404.90 m. sample: cs, p39604, 1. optical parameters: d, 153 × 1034, p46/2‒4, × 40 pc. repository: gsc 138034. fig. 5. cicatricososporites eocenicus. location: south labrador n-79, 2040‒2050 m. sample: cs, p39786, 1. optical parameters: c, 187 × 1040, t32/1–t33/2, × 50 bf. repository: gsc 138118. fig. 6. baculatisporites crenulatus sp. nov. location: roberval k-92, 2090‒2100 m. sample: cs, p17692, 1. optical parameters: c, 45 × 1091, d39/0, × 50 bf. repository: gsc 137890. fig. 7. baculatisporites crenulatus sp. nov. location: roberval k-92, 2150‒2160 m. sample: cs, p17696, 1. optical parameters: c, 63 × 1073, f37/0, × 50 bf. repository: gsc 137891. fig. 8. baculatisporites crenulatus sp. nov., holotype. location: roberval k-92, 1790‒1800 m. sample: cs, p17684, 1. optical parameters: c, 106 × 1048, l34/2, × 50 bf. repository: gsc 137889. fig. 9. caryapollenites inelegans. location: karlsefni a-13, 969.28–978.42 m. sample: cs, p39553, 1. optical parameters: d, 81 × 1060, h49/2, × 40 pc. repository: gsc 138003. fig. 10. caryapollenites inelegans. location: rut h-11, 845‒855 m. sample: cs, p39322, 1. optical parameters: d, 112 × 844, l27/1, × 40 pc. repository: gsc 137925. fig. 11. caryapollenites veripites. location: karlsefni a-13, 2231.16–2240.31 m. sample: cs, p39598, 1. optical parameters: d, 191 × 993, t42/0, × 40 pc. repository: gsc 138028. fig. 12. caryapollenites veripites. location: karlsefni a-13, 3843.57–3852.72 m. sample: cs, p39656, 1. optical parameters: d, 164 × 867, q29/0, × 40 pc. repository: gsc 138056. fig. 13. momipites annellus. location: karlsefni a-13, 3904.54–3913.68 m. sample: cs, p39658, 1. optical parameters: d, 175 × 911, r34/3, × 40 pc. repository: gsc 138058. fig. 14. momipites annellus. location: karlsefni a-13, 3301.02–3310.17 m. sample: cs, p39637, 1. optical parameters: d, 122 × 917, m34/0, × 40 pc. repository: gsc 138051. fig. 15. momipites coryloides. location: snorri j-90, 1033.28‒1043.43 m. sample: cs, p9709, 10. optical parameters: c, 103 × 1078, k37–k38, × 50 bf. repository: gsc 138133. fig. 16. callialasporites dampieri. location: karlsefni a-13, 969.28–978.42 m. sample: cs, p39553, 1. optical parameters: d, 196 × 1026, t45/4, × 40 pc. repository: gsc 138006. fig. 17. chenopodipollis sp. location: karlsefni a-13, 911.36–920.51 m. sample: cs, p39551, 1. optical parameters: d, 115 × 930, l35/2‒4, × 40 pc. repository: gsc 138001. fig. 18. compositoipollenites sp. b. of williams & brideaux 1975. location: karlsefni a-13, 1435.63–1444.77 m. sample: cs, p39570, 1. optical parameters: d, 187 × 934, s36/3, × 40 pc. repository: gsc 138012. fig. 19. corsinipollenites oculusnoctis. location: south labrador n-79, 1110‒1120 m. sample: cs, p39757, 1. optical parameters: c, 200 × 1018, u31/1–4, × 50 bf. repository: gsc 138080. fig. 20. corsinipollenites oculusnoctis. location: karlsefni a-13, 1709.95–1719.09 m. sample: cs, p39579, 1. optical parameters: d, 196 × 1042, t47/3, × 40 pc. repository: gsc 138014. plate 18 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 138 139 1 9 5 13 17 2 10 6 14 18 3 11 7 15 19 4 12 8 16 20 plate 18 plate 18 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 139 140 fig. 1. cicatricosisporites ornatus. location: rut h-11, 2315‒2325 m. sample: cs, p39371, 1. optical parameters: d, 163 × 885, q31/0, × 40 pc. repository: gsc 137949. fig. 2. cicatricosisporites ornatus. location: rut h-11, 1865‒1875 m. sample: cs, p39356, 1. optical parameters: d, 115 × 881, l30/2‒4, × 40 pc. repository: gsc 137942. fig. 3. cicatricosisporites ornatus. location: karlsefni a-13, 2532.98–2542.08 m. sample: cs, p39609, 1. optical parameters: d, 163 × 873, q30/1‒3, × 40 pc. repository: gsc 138038. fig. 4. cicatricosisporites ornatus. location: rut h-11, 3695‒3705 m. sample: cs, p39416, 1. optical parameters: d, 90 × 993, h42/3‒4, × 40 pc. repository: gsc 137970. fig. 5. extratriporopollenites sp. location: gilbert f-53, 1540‒1550 m. sample: cs, p39464, 1. optical parameters: c, 180 × 1043, s34/0, × 50 bf. repository: gsc 137975. fig. 6. extratriporopollenites sp. location: hekja o-71, 1610‒1620 m. sample: cs, p20568, 1. optical parameters: c, 197 × 1032, u33/1, × 50 bf. repository: gsc 137914. fig. 7. extratriporopollenites sp. location: hekja o-71, 1420‒1430 m. sample: cs, p20562, 1. optical parameters: c, 80 × 1077, h37/2, × 50 bf. repository: gsc 137913. fig. 8. extratriporopollenites sp. location: snorri j-90, 1892.83‒1901.98 m. sample: cs, p9737, 1. optical parameters: c, 135 × 1020, n31/4–n32/3, × 50 bf. repository: gsc 138135. fig. 9. extratriporopollenites sp. location: south labrador n-79, 780‒790 m. sample: cs, p39746, 1. optical parameters: c, 154 × 1018, p31/4, × 50 bf. repository: gsc 138078. fig. 10. periporopollenites sp. location: karlsefni a-13, 2176.30–2185.44 m. sample: cs, p39596, 1. optical parameters: d, 109 × 957, k38/4, × 40 pc. repository: gsc 138023. fig. 11. periporopollenites sp. location: rut h-11, 1305–1315 m. sample: cs, p39337, 1. optical parameters: d, 148 × 859, o28/0, × 40 pc. repository: gsc 137935. fig. 12. graminidites sp. a of williams & brideaux 1975. location: rut h-11, 1685‒1695 m. sample: cs, p39350, 1. optical parameters: d, 177 × 882, r31/3, × 40 pc. repository: gsc 137937. fig. 13. pistillipollenites macgregorii. location: snorri j-90, 2441.48‒2450.62 m. sample: cs, p9751, 10. optical parameters: c, 30 × 910, c20/0, × 50 bf. repository: gsc 138140. fig. 14. pistillipollenites macgregorii. location: gjoa o-37, 1620‒1630 m. sample: cs, p16951, 1. optical parameters: c, 65 × 1174, f37/4, × 50 bf. repository: gsc 137885. fig. 15. quercoidites sp. location: rut h-11, 1205‒1215 m. sample: cs, p39334, 1. optical parameters: d, 119 × 869, l29/4, × 40 pc. repository: gsc 137930. fig. 16. quercoidites sp. location: rut h-11, 1305‒1315 m. sample: cs, p39337, 1. optical parameters: d, 102 × 957, k38/2, × 40 pc. repository: gsc 137934. fig. 17. tiliaepollenites crassipites. location: south labrador n-79, 1110‒1120 m. sample: cs, p39757, 1. optical parameters: c, 182 × 1009, s30/2, × 50 bf. repository: gsc 138079. fig. 18. tiliaepollenites crassipites. location: south labrador n-79, 780‒790 m. sample: cs, p39746, 1. optical parameters: c, 134 × 1023, n32/3, × 50 bf. repository: gsc 138077. fig. 19. tiliaepollenites sp. a. location: rut h-11, 1565‒1575 m. sample: cs, p39346, 1. optical parameters: d, 190 × 928, t35/2, × 40 pc. repository: gsc 137936. fig. 20. tiliaepollenites sp. a. location: karlsefni a-13, 911.36–920.51 m. sample: cs, p39551, 1. optical parameters: d, 134 × 1000, n43/0, × 40 pc. repository: gsc 138002. plate 19 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 140 141 1 9 5 13 17 2 10 6 14 18 3 11 7 15 19 4 12 8 16 20 plate 19 plate 19 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 141 142 fig. 1. osmundacidites wellmannii. location: rut h-11, 905‒915 m. sample: cs, p39324, 1. optical parameters: d, 191 × 968, t39/2, × 40 pc. repository: gsc 137927. fig. 2. rugubivesiculites sp. location: ogmund e-72, 1965 m. sample: cs, yd15763, 3. optical parameters: a, 35.4 × 103.2, n35/2, × 60 bf. repository: mguh 31361. fig. 3. wodehouseia spinata. location: skolp e-07, 1015 m. sample: cs, yd15576, 3. optical parameters: a, 52.0 × 100.5, q53/1, × 60 bf. repository: mguh 31362. fig. 4. zlivisporis sp. location: karlsefni a-13, 1792.25–1801.39 m. sample: cs, p39582, 1. optical parameters: d, 96 × 851, j27/2, × 40 pc. repository: gsc 138017. fig. 5. zlivisporis sp. location: rut h-11, 2285‒2295 m. sample: cs, p39370, 1. optical parameters: d, 133 × 1041, n47/0, × 40 pc. repository: gsc 137948. fig. 6. zlivisporis sp. location: karlsefni a-13, 2148.87–2158.01 m. sample: cs, p39595, 1. optical parameters: d, 110 × 1047, k48/3, × 40 pc. repository: gsc 138022. fig. 7. zlivisporis sp. location: rut h-11, 1805‒1815 m. sample: cs, p39354, 1. optical parameters: d, 120 × 1031, m46/1, × 40 pc. repository: gsc 137940. fig. 8. zlivisporis sp. location: ogmund e-72, 1785 m. sample: cs, yd15751, 3. optical parameters: a, 17.5 × 111.0, e17/2, × 60 bf. repository: mguh 31363. fig. 9. zonalapollenites igniculus. location: karlsefni a-13, 859.55–865.64 m. sample: cs, p39549, 1. optical parameters: d, 100 × 954, j38/3‒4, × 40 pc. repository: gsc 138000. fig. 10. zonalapollenites igniculus. location: rut h-11, 845‒855 m. sample: cs, p39322, 1. optical parameters: d, 179 × 914, s34/1, × 40 pc. repository: gsc 137926. fig. 11. zonalapollenites igniculus. location: south labrador n-79, 720‒730 m. sample: cs, p39744, 1. optical parameters: c, 142 × 1009, o30/4–p30/2, × 50 bf. repository: gsc 138076. fig. 12. fungal element: callimothallus. location: karlsefni a-13, 2478.05–2487.20 m. sample: cs, p39607, 1. optical parameters: d, 161 × 1054, q48/0, × 40 pc. repository: gsc 138035. fig. 13. fungal element: diporicellaesporites. location: karlsefni a-13, 2505.49–2514.63 m. sample: cs, p39608, 1. optical parameters: d, 127 × 964, m39/0, × 40 pc. repository: gsc 138036. fig. 14. fungal element: fractisporonites. location: karlsefni a-13, 1737.38–1746.53 m. sample: cs, p39580, 1. optical parameters: d, 179 × 889, r31/4, × 40 pc. repository: gsc 138015. fig. 15. fungal element: fusiformisporites. location: rut h-11, 2285‒2295 m. sample: cs, p39370, 1. optical parameters: d, 106 × 906, k33/0, × 40 pc. repository: gsc 137946. fig. 16. fungal element: microthallites. location: karlsefni a-13, 1764.81–1773.96 m. sample: cs, p39581, 1. optical parameters: d, 108 × 887, k31/0, × 40 pc. repository: gsc 138016. fig. 17. fungal element: staphlosporonites. location: karlsefni a-13, 2286.03–2295.17 m. sample: cs, p39600, 1. optical parameters: d, 141 × 1071, o50/0, × 40 pc. repository: gsc 138031. fig. 18. fungal element: multicellaesporites. location: karlsefni a-13, 1792.25–1801.39 m. sample: cs, p39582, 1. optical parameters: d, 186 × 885, s31/0, × 40 pc. repository: gsc 138018. fig. 19. fungal element: multicellaesporites. location: rut h-11, 3725‒3735 m. sample: cs, p39417, 1. optical parameters: d, 175 × 1043, r47/4, × 40 pc. repository: gsc 137972. fig. 20. fungal element: pluricellaesporites. location: rut h-11, 2435‒2445 m. sample: cs, p39375, 1. optical parameters: d, 160 × 887, q31/1‒2, × 40 pc. repository: gsc 137953. plate 20 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 142 143 1 9 5 13 17 2 10 6 14 18 3 11 7 15 19 4 12 8 16 20 plate 20 plate 20 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 143 bulletin36.qxp_bulletin 36 19/12/16 13.40 side 144 << /ascii85encodepages false /allowtransparency true /autopositionepsfiles true /autorotatepages /none /binding /left /calgrayprofile (dot gain 20%) 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/expandpage false /honorbaseurl true /honorrollovereffect false /ignorehtmlpagebreaks false /includeheaderfooter false /marginoffset [ 0 0 0 0 ] /metadataauthor () /metadatakeywords () /metadatasubject () /metadatatitle () /metricpagesize [ 0 0 ] /metricunit /inch /mobilecompatible 0 /namespace [ (adobe) (golive) (8.0) ] /openzoomtohtmlfontsize false /pageorientation /portrait /removebackground false /shrinkcontent true /treatcolorsas /mainmonitorcolors /useembeddedprofiles false /usehtmltitleasmetadata true >> ] >> setdistillerparams << /hwresolution [2400 2400] /pagesize [612.000 792.000] >> setpagedevice geological survey of denmark and greenland bulletin 4, pp 1-7 geological survey of denmark and greenland bulletin 4 • 2004 review of survey activities 2003 edited by martin sønderholm and a.k. higgins geological survey of denmark and greenland ministry of the environment geological survey of denmark and greenland bulletin 4 keywords geological survey of denmark and greenland, survey organisations, current research, denmark, greenland, faroe islands. cover photographs from left to right 1. field work in west greenland. photo: jakob lautrup. 2. petroleum geological field work in the song ba trough, a cenozoic rift-lake basin in central vietnam (see article on page 97). photo: henrik ingermann petersen. 3. fossil hunting following the major rock fall at store stejlebjerg, møn, denmark in 2003 (see article on page 89). photo: peter k. warna-moors. 4. installation of an automatic mass balance station on the greenland inland ice. the station carries out ablation measurements that are transmitted by satellite every 6 hours to the survey (see article on page 81). photo: carl egede bøggild. frontispiece: facing page preparation of field equipment for a sampling and ground magnetic survey of a large kimberlitic dyke in the maniisoq region, southern west greenland. approximately 1.1 tonnes of kimberlitic rock was collected along the 2.5 km long dyke and subsequently tested for diamond content by a canadian laboratory (see article on page 69). photo: jakob lautrup. chief editor of this series: peter r. dawes scientific editors: martin sønderholm and a.k. higgins editorial secretaries: esben w. glendal and birgit eriksen illustrations: jette halskov lay-out and graphic production: annabeth andersen and henrik klinge pedersen printers: schultz grafisk, albertslund, denmark manuscripts submitted: 3 december 2003 – 23 march 2004 final version approved: 12 may 2004 printed: 20 july 2004 isbn 87-7871-132-0 issn 1603-9769 geological survey of denmark and greenland bulletin the series geological survey of denmark and greenland bulletin replaces geology of denmark survey bulletin and geology of greenland survey bulletin. citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 4, 100 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk or geografforlaget aps rugårdsvej 55, dk-5000 odense c, denmark phone: +45 63 44 16 83, fax: +45 63 44 16 97, e-mail: go@geografforlaget.dk  danmarks og grønlands geologiske undersøgelse (geus), 2004 4 review of survey activities k. sørensen 7 the history of hydrocarbon filling of danish chalk fields p. frykman, o.v. vejbæk, n. bech and c.m. nielsen 9 assessing the european potential for geological storage of co2: the gestco project n.p. christensen and m. larsen 13 geothermal energy in denmark l.h. nielsen, a. mathiesen and t. bidstrup 17 the billund delta: a possible new giant aquifer in central and western jutland e.s. rasmussen, k. dybkjær and s. piasecki 21 pesticide leaching in danish groundwater: identification of vulnerable areas e. nygaard, v. ernstsen, c.s. jacobsen, o.h. jacobsen, r.k. juhler, p. van der keur, s.e. olesen, j. rasmussen, p. rosenberg and h. vosgerau 25 immunological analysis of pesticides: a new tool in groundwater testing j. aamand, l. bruun and c.b.v. christensen 29 direct analysis of microbial populations in soil and freshwater aquifers by using nucleic acid based techniques c.s. jacobsen, j.r. de lipthay, m. bender, l. fredslund, a.r. johnsen and k. johnsen 33 using the geological record to assess the changing status of danish lakes e.g. bradshaw and p. rasmussen 37 sediment distribution and transport in the shallow coastal waters along the west coast of denmark j.o. leth, b. larsen and d. anthony 41 ghana burkina faso dominican republic uganda tanzania mozambique latvia ukraine romania denmark faroe islands greenland contents geus working areas 2003. orange areas are covered in this volume. 5 environmental data and the internet: openness and digital data management j. tulstrup 45 age of oils in west greenland: was there a mesozoic seaway between greenland and canada? j.a. bojesen-koefoed, h.p. nytoft and f.g. christiansen 49 seismic and petrophysical properties of faroe islands basalts: the seifaba project p. japsen, m.s. andersen, l.o. boldreel, r. waagstein, r.s. white and m. worthington 53 geohazard studies offshore the faroe islands: slope instability, bottom currents and sub-seabed sediment mobilisation t. nielsen and a. kuijpers 57 exploring for extended continental shelf claims off greenland and the faroe islands – geological perspectives c. marcussen, f.g. christiansen, t. dahljensen, m. heinesen, s. lomholt, j.j. møller and k. sørensen 61 gold in central west greenland – known and prospective occurrences a. steenfelt, h. stendal, b.m. nielsen and t.m. rasmussen 65 investigating the diamond potential of southern west greenland s.m. jensen and k. secher 69 low-pressure metamorphism during archaean crustal growth: a low-strain zone in the northern nagssugtoqidian orogen, west greenland a.a. garde, m.s. christiansen, j.a. hollis, s. mazur and j.a.m. van gool 73 epithermal gold and massive sulphide mineralisation in oil impregnated palaeogene volcanic rocks of ubekendt ejland, west greenland s. bernstein and c. knudsen 77 towards an assessment of the balance state of the greenland ice sheet c.e. bøggild, c. mayer, s. podlech, a. taurisano and s. nielsen 81 jakobshavn isbræ, west greenland: the 2002–2003 collapse and nomination for the unesco world heritage list a. weidick, n. mikkelsen, c. mayer and s. podlech. 85 prediction and risk evaluation of chalk cliff collapse: the protect project s.a.s. pedersen and i. møller 89 co-operation with the geological survey department of ghana f. kalsbeek, b. hermansen, c. knudsen, l. thorning and m. thorsen 93 petroleum potential of sedimentary basins in vietnam: long-term geoscientific co-operation with the vietnam petroleum institute l.h. nielsen and i. abatzis 97 lao pdr vietnam mongolia kyrgyz republic 7 one of the last visible relics of the pre-fusion history of the geological survey of denmark and greenland (geus), the two separate bulletin series for greenland and denmark, respectively, ceased to appear in late 2003. almost a decade after the merging of the geological survey of denmark (dgu) and the geological survey of greenland (ggu), the two bulletin series were merged to form the geological survey of denmark and greenland bulletin, the first issue of which is the monumental the jurassic of denmark and greenland. with the closure of the two old series, a long-standing tradition of publishing an annual review of greenland activities was brought to an end. everyone in geus, not least those who grew up within the former dgu, agree that the review of greenland activities was a high-quality, enjoyable-to-read review of the new survey’s greenland activities. it was therefore a natural progression to publish a review volume covering the full spectrum of activities of the entire institution, a review of survey activities intended for a professional, but not specialist, readership. in order to keep the volume to a manageable size, all articles have been restricted to a four-page limit. thanks to the diverse geology of denmark and greenland, and the reliance of society on resources hosted in the subsurface, geus as a survey has been generously endowed by nature. however, as for many other geological surveys in western europe direct financial government appropriations for geus have decreased significantly during the past several years. increasing funding from external scientific foundations and commercial sources has proved to be a partial compensation. furthermore, the survey has extended the traditional scope of its activities to also include capacity-building projects in several developing countries within the fields of institutional development, geological mapping, petroleum geology, mineral exploration and hydrogeology. despite recent hardships, including an inevitable loss of personnel, the size, relevance and quality of the geological output of the survey are still undergoing a healthy development. the articles contained in this bulletin review many of the principal activities of the survey in 2003; they reflect the diversity of our survey, from the microbial to the plate tectonic level. of the 23 articles, 20 measure directly the extent to which the survey’s activities have been driven by external demand, inasmuch as they describe activities which have been undertaken for external customers or rely on external sources for a significant funding contribution. funding for these 20 projects has come from national and international research funds, from the petroleum industry, from ministries and counties, from the bureau of minerals and petroleum of the greenland home rule government, and from national and international funds for aid to developing countries. it is expected that the focus on applied geological projects will also be reflected in future volumes of review of survey activities. review of survey activities kai sørensen director geological survey of denmark and greenland bulletin 23, 2011, 65–68 65 this paper describes structural data collected during field work in southern east greenland, a region characterised by a complex tectonic history. here, 3d photogeology based on aerial and oblique photographs using high-resolution photogrammetry of a 150 km2 area in sødalen in southern east greenland shows ese–wnw-trending faults cross-cutting paleocene rift structures and flexure-related normal faults. the kinematic analysis highlights oblique and left-lateral strike-slip movements along faults oriented 120°. strike-slip and dip-slip kinematic indicators on the walls of the chilled contacts between alkaline e–w-oriented dykes and the volcanic host rocks suggest that the faults and dykes formed at the same time, or maybe the faults were re-activated at a later stage. palaeostress analysis, performed by inversion of fault-slip data, shows the presence of three different tectonic events. coupling the 3d photogeological tool with structural analysis at key localities is a fundamental way to understand better the tectonic history of such a large area. geological setting the blosseville kyst in southern east greenland is characterised by a thick sequence of flood basalts and mafic intrusions (fig. 1). the skaergaard layered gabbro, the miki fjord macrodyke and dolerite sill complexes were formed during the continental break-up and the initial opening of the north–east atlantic ocean at 55 ma (nielsen 1975; karson & brooks 1999; tegner et al. 2008). in the sødalen region pre-basaltic sediments characterise the kangerlussuaq basin and the lower part of the blosseville group (wager 1947; nielsen et al. 1981). sedimentological studies recognise different facies associations of late aptian to late paleocene age (larsen, m. et al. 1999). the youngest part of the basin comprises interfingering paleocene volcanic units. based on stratigraphy, geochemistry and petrography, the lavas of the blosseville group have been divided into two main series: (1) a 2 km thick sequence of volcanic rocks that formed in a continental rift environment (nielsen et al. 1981), and (2) a 6 km thick sequence of plateau basalts (larsen, l.m. et al. 1989). furthermore, the blosseville kyst is characterised by different generations of dykes and sills, partly related to the break-up and post-break-up history (wager 1947; hanghøj et al. 2003). southern east greenland is a type example of a volcanic rifted margin (geoffroy 2005). the geological evolution of the margin is interpreted as the result of a ne–sw-oriented late cretaceous rifting phase that led to the onset of oceanic spreading in the late paleocene – early eocene (c. 55 ma) after a period of syn-rift continental tectonism and volcanism. the general south-east dip of the basalts, the presence of landwarddipping normal faults and the coastal dyke swarm suggest a regional lithosphere flexure (larsen, h.c. & saunders 1998). sødalen region sødalen is an 8 km long, nw–se-oriented, u-shaped glacial valley extending se–nw up to the ‘sødalengletscher’ (fig. 2a). the bedrock of the area is characterised by gneiss basement, locally overlain by syn-rift sedimentary and volcanic rocks that form a monocline that dips south-eastwards. the late paleocene syn-rift sedimentary rocks crop out along the western side of the valley; they are unconformably overlain by sedimentary rocks belonging to the vandfaldsdalen formaanalysis of palaeogene strike-slip tectonics along the southern east greenland margin (sødalen area) pierpaolo guarnieri sø dalen ‘sødalengletscher’ 68°15´n 31°w miki fjord skaergaard intrusion m ac ro dy ke fig. 2a greenland mainly palaeogene volcanic rocks palaeogene gabbro mesozoic–palaeogene sedimentary rocks precambrian basement 10 km b l o s s e v i l l e k y s t fig. 1. simplified geological map of the sødalen region in southern east greenland. © geus, 2011. geological survey of denmark and greenland bulletin 23, 65–68. open access: www.geus.dk/publications/bull 6666 main strike-slip fault faults and dykes bedding structural locality 1 10 sill miki fjord macrodyke mikis fm lava flows breccias hyaloclastite volcaniclastics schjelderup member precambrian basement vandfaldsdalen formation 300 2 0 0 400 50 0 100 700 800 6 0 0 700 600 50 0 500 300 500 900 60 0 60 0 800 800 400 4 0 0 80 070 0 300 200 900 8 7 4 8 5 9 7 19 20 22 10 15 10 900 sødalen 1 km 68°14´n 68°12´n 68°14´n 68°12´n 31°25´w 31°20´w 31°20´w31°25´w a b c d lake ice cap quaternary deposits ‘s ød al en gl et sc he r’ 5 2 4 3 1 faults (outcrop data) dykesfaultsfaults and dykes (vertical photographs) reverse faultsnormal faults left-lateral faults all faults (n = 92) n nnn s1> s2> n = 111 max = 13% n = 66 max = 18% n = 82 max = 17% n = 92 max = 15% (oblique photographs) right-lateral faults s3 n nn phase 1 phase 2 phase 3 67 tion (nielsen et al. 1981). the unconformity may be related to pre-volcanic uplift, coeval with the ne–sw-oriented rifting, followed by a rapid subsidence that accommodated the volcanism (larsen, m. et al. 1999). the continental breakup is contemporaneous with the emplacement of layered gabbro bodies dated to c. 55 ma, which formed at c. 2 km depth in the continental crust. the skaergaard intrusion and the miki fjord macrodyke (nielsen et al. 1981; tegner et al. 2008) are contemporaneous with the up to 6 km thick sequence of plateau basalts (larsen, l.m. et al. 1989). structural data a total of 350 measurements for structural analysis were collected, from two sources: (1) from outcrops (metre scale) at five sites used for kinematic analysis and (2) from 3d photo geology to evaluate strike and dip direction and cross-cutting relationships of faults and dykes using vertical aerial photographs (kilometre scale) and oblique photographs (100 m scale). a new tool for photogeology and mapping is developed and implemented at geus to collect geological features as 3d polylines with a descriptive gis database suitable for 3d modelling (vosgerau et al. 2010). dykes – three main generations of dykes are found in the area. their relative ages can be established from cross-cutting relationships, which show that the oldest generation (d1) is mainly ne–sw-oriented, orthogonal to bedding or landward-dipping; the trend is parallel to the miki fjord macrodyke. the average trend of the second generation (d2) is ene–wsw and these dykes are almost vertical (fig. 2b). the third generation of dykes found in the area (d3) trends e–w (fig. 2b). faults – two main trends of fault traces, up to 2 km long, can be followed on the vertical aerial photographs (fig. 2b). the oldest generation (f1) is characterised by ene–wsworiented normal faults. these faults are mainly landwarddipping and are interpreted as flexure-related faults (wager 1947; nielsen et al. 1981). at site 3 (fig. 2a), the miki fjord macrodyke contact is downfaulted by a landward-dipping (f1) normal fault with an average vertical offset of 400 m. the youngest (f2) faults trend ese–wnw. south-east of localities 2 and 4 (fig. 2a), the fault traces are curved in planar view typical of strike-slip fault systems. kinematic analysis field data suitable for fault-slip analysis include measurements of fault plane orientations, slip directions, senses of slip and bedding orientations. the slip direction of faults is determined using slickensides and calcite fibres on the fault plane. sense of slip indicators include tails and scratches and crescentic marks formed by intersection of the fault plane with secondary fractures such as: r, r’, p and t (petit 1987). data collected in the canyon at locality 2 (fig. 2a) define the kinematics of a 120°-trending fault corresponding to a major left-lateral strike-slip fault that cuts the basalts. the fault zone is c. 50 m wide and contains a >50 cm thick calcite vein. double movement along the fault plane with well developed dip-slip and strike-slip slickensides and calcite fibres suggests a reactivation of the fault (fig. 3). the estimated vertical offset, based on the tectonic contact between two stratigraphic markers, is around 250 m, whereas the horizontal offset is estimated to 500 m. this results in a more than 50–150 m wide, 120°-trending, rhomb-shaped fault zone, 1 km long in map view (fig. 2a, locality 2) and with a negative flower structure in cross-section. to the south-east, the fault trace disappears below an ice cap and to the north-west it is covered by the moraine in front of ‘sødalengletscher’, but it is exposed on the western side of sødalen, where a well facing page: fig. 2. structural data analysis. a: geological map of the sødalen area (modified from nielsen et al. 1981). arrows show the direction of movement along strike-slip faults; contour lines 100 m. b: rose diagrams for orientation of faults and dykes. c: lower hemisphere stereographic projection of faults grouped by kinematics; arrows show the slip vector. d: palaeostress analysis of 92 fault-slip measurements. black arrows indicate maximum horizontal shortening/extension; σ1, σ2, σ3 = principal axes of stress. visualisation of the right dihedral method (red = pressure, blue = tension) shows planes that are likely to have been re-activated (the three diagrams to the right). dip slipdip slip strike-slipstrike-slipstrike-slip dip slip fig. 3. evidence of multiple re-activation of a fault testified by well-developed dip-slip and strike-slip slickensides on a fault plane (locality 2 in fig. 2a). 6868 developed vertical cleavage, locally with strike-slip slickensides, cross-cuts the miki fjord macrodyke. at locality 5 (fig. 2a), a 4 km long e–w-oriented dyke crosses sødalen; it is an example of the latest dyke generation (d3). slickensides are found on the chilled margins of the dyke, which show that both dip-slip and strike-slip movements have taken place. the trend of the dykes, coupled with evidence of multiple reactivation of the contact, suggests a relationship between strike-slip faults and dykes in which normal faults intruded by dykes were re-activated as leftlateral faults in a nne–ssw extensional regime associated with the ese–wnw-trending shear-zone (fig. 2a). palaeostress analysis palaeostress analysis of the heterogeneous fault-slip data set was performed using integrated software for structural analysis (žalohar 2009). more than 90 fault-slip measurements were taken at five sites (localities 2a–c) and used for inversion to obtain palaeostress values. the gauss method associated with visualisation of p&t dihedra (žalohar 2009) distinguishes three superimposed tectonic phases in the area (fig. 2d): (1) a phase with strike-slip regime and a 20–30° trending maximum horizontal shortening interpreted as oblique rifting; (2) a phase with a sse–nnw-trending maximum horizontal extension that corresponds to the coastal flexure and (3) a phase with strike-slip regime and a 95°-trending maximum horizontal shortening that caused the inversion and uplift of the entire area. conclusions the structural data collected in sødalen indicate the presence of strike-slip faults related to two tectonic events separated in time by the coastal flexure. the youngest structures and dykes (phase 3; fig. 2d) are associated with a nw–se left-lateral shear zone that cross-cuts the paleocene rift and the structures related to the coastal flexure of the continental margin (phase 2). the evidence of dyke intrusions related to n–s extension compatible with the strike-slip tectonic regime of phase 3, suggests a coexistence of the two phenomena as a superficial expression of deep-seated crustal structures. the oldest structures and dykes of phase 1 show a maximum horizontal extension coherent with the trend of the miki fjord macrodyke. this strike-slip tectonic regime could be related to an oblique rifting stage in paleocene time. finally, the accuracy of the 3d photogeological tool is tested over a range of kilometre to metre scale (fig. 2b) showing the power of this method developed at the survey. acknowledgement chevron is thanked for interest and financial support. references geoffroy, l.: 2005. volcanic passive margins. comptes rendus geoscience 337, 1395–1408. hanghøj k., storey m. & stecher o. 2003: an isotope and trace element study of the east greenland tertiary dyke swarm: constraints on temporal and spatial evolution during continental rifting. journal of petrology 44, 2081–2112. karson, j.a. & brooks, c.k. 1999: structural and magmatic segmentation of the tertiary east greenland volcanic rifted margin. in: ryan, p.d. (ed.): continental tectonics. geological society special publications (london) 164, 313–338. larsen, h.c. & saunders a.d. 1998: tectonism and volcanism at the southeast greenland rifted margin: a record of plume impact and later continental rupture. proceedings of the ocean drilling program, scientific results 152, 503–534. larsen, l.m., watt, w.s. & watt, m. 1989: geology and petrology of the lower tertiary plateau basalts of the scoresby sund region, east greenland. bulletin grønlands geologiske undersøgelse 157, 164 pp. larsen, m., hamberg, l., olaussen, s., nørgaard-pedersen, n. & stemmerik, l. 1999: basin evolution in southern east greenland: an outcrop analog for cretaceous–paleogene basins on the north atlantic volcanic margins. aapg bulletin 83, 1236–1261. nielsen,t.f.d. 1975: possible mechanism of continental breakup in the north atlantic. nature 253, 182–184. nielsen, t.f.d., soper, n.j., brooks, c.k., faller, a.m., higgins, a.c. & matthews, d.w. 1981: the pre-basaltic sediments and the lower basalts at kangerdlugssuaq, east greenland: their stratigraphy, lithology, palaeomagnetism and petrology. meddelelser om grønland geoscience 6, 3–25. petit, j.p. 1987: criteria for the sense of movement on fault surfaces in brittle rocks. journal of structural geology 9, 597–608. tegner, c., brooks, c.k., duncan, r.a., heister, l.e. & bernstein, s. 2008: 40ar-39ar ages of intrusions in east greenland: rift-to-drift transition over the iceland hotspot. lithos 101, 480–500. vosgerau h., guarnieri p., weibel r., larsen m., dennehy, c., sørensen, e.v. & knudsen, c. 2010: study of a palaeogene intrabasaltic sedimentary unit in southern east greenland: from 3-d photogeology to micropetrography. geological survey of denmark and greenland bulletin 20, 75–78. wager, l.r. 1947: geological investigations in east greenland, part iv: the stratigraphy and tectonics of knud rasmussens land and the kangerdlugssuaq region. meddelelser om grønland 134(5), 62 pp. žalohar, j. 2009: t-tecto 3.0 professional. integrated software for structural analysis of fault-slip data. department of geology, si-1000 ljubljana, slovenia. author’s address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: pgua@geus.dk geological survey of denmark and greenland bulletin 26, 2012, 37-40 37 shallow geothermal energy in denmark thomas vangkilde-pedersen, claus ditlefsen and anker lajer højberg the use of shallow geothermal energy instead of fossil fuels can lead to substantial reductions in co2 emissions. however, the use of shallow geothermal energy in denmark is limited compared to, e.g. sweden and germany and we still lack know-how and experience with its use in denmark. in co-operation with research and industry partners, the geological survey of denmark and greenland is conducting a three-year project geoenergy, tools for ground-source heating and cooling based on closed-loop boreholes (www.geoenergi.org). the objective of the project is to acquire knowledge and develop tools and best practice for the design and installation of shallow geothermal energy systems. exploitation of shallow geothermal energy in the shallow subsurface, solar energy absorbed and stored in the ground determines the temperature, whereas the temperature gradient is determined by the heat flux from the interior of the earth. in denmark the net insolation is c. 400 kwh/m2/year and in the upper few hundred metres, the heat flux from the interior is c. 0.20 to 0.35 kwh/m2/year (balling et al. 1992). the geothermal gradient in denmark is 25–30°c per kilometre (mathiesen et al. 2009) with an upper zone of seasonal variations. the thickness of this zone has not been investigated in detail in denmark, but a general thickness of 10–20 m may be expected (fig. 1; banks 2008). the shallow geothermal energy is exploited in combination with heat pumps bringing the temperature to the desired operating temperature of the heating (or cooling) system. a ground-source heat pump, however, requires energy in the form of heat or cold from the ground as well as electricity for the operation of the pump itself. the energy extraction from the ground can be based on either open-loop systems or closed-loop systems. in open-loop systems, groundwater from a production well is used directly as an energy source for the heat pump and returned to the aquifer via an injection well. the efficiency is high due to a constant and relatively high temperature of the water (typically 8–10°c), but a suitable aquifer with sufficient yield is required and conflicts of interest with neighbouring installations or nearby drinking water catchment may arise. in closed-loop systems, water with antifreeze is circulated in the ground in high-density polyethylene pipes and acts as a heat exchanger. the heat capacity of soil and groundwater is relatively high, whereas the thermal conductivity is moderate and dependent on the local hydrogeology. therefore the heat exchanger must be relatively large, i.e., the polyethylene pipes relatively long. they can either be installed horizontally at a depth of c. 1 m (horizontal closed-loop systems), or vertically in a borehole (borehole heat exchanger or vertical closed-loop system; fig. 2). vertical systems are only influenced by seasonal temperature variations in the upper zone down to 10–20 m and have a higher efficiency than horizontal systems due to a higher and constant ground temperature. on the other hand, the temperature regime around a vertical system is only slowly re-generated and careful design with respect to the energy demand is necessary, and an accurate estimate of the energy consumption is needed to optimise the design of the system. alternating operations with heating during winter and cooling during summer, supplemented with seasonal storage of fig. 1. schematic block diagram showing the geothermal gradient and magnitude of geothermal heat flux and insolation with average values for denmark as well as the seasonal zone of fluctuation in temperature (modified from banks 2008). atmospheric longwave radiation 925–1000 kwh m –2 year –1 back radiation insolation changes winter summer geothermal heat flux temperature 0.20–0.35 kwh m –2 year –1 heat transfer via conduction, convection and evapotranspiration g eo therm al gradient c. 25–30° km –1 m ean annual tem perature c. 400 kwh m –2 year –1 net radiation seasonal © 2012 geus. geological survey of denmark and greenland bulletin 26, 37–40. open access: www.geus.dk/publications/bull 3838 heat from, e.g. solar panels, can optimise the efficiency of vertical systems. groundwater and environmental protection protection of the environment and groundwater is of paramount importance in shallow geothermal projects. the main issues are leakage of water with antifreeze, cross-connecting different aquifers, seepage of surface water along the borehole, drilling into artesian aquifers and unwanted thermal effects. all steps in the production and installation of a ground loop must be designed for optimum tightness and longevity of the loop. installation procedures must minimise the risk of damaging the pipes, and a pressure test of the pipes must be carried out. the legislation also specifies which antifreeze agents can be used in the ground loop – only non-toxic and easily biodegradable fluids are allowed. closed-loop boreholes must be sealed with low permeable grout in order to prevent contact between the surface and aquifers or between individual aquifers and the quality of the sealing determines the protection of the groundwater, as for all other boreholes. grouting is also necessary to improve the heat transfer from the ground to the pipes and in actual practice the entire borehole is sealed with grout. perforation of an artesian aquifer can cause large quantities of water to press through the drill pipe disrupting the grout. to avoid drilling closed-loop boreholes into artesian aquifers, local hydraulic heads must be checked prior to drilling (ditlefsen 2012). possible problems with thermal pollution for horizontal systems are described by banks (2008). extraction of heat may cause freezing of the ground and frost heaving, which can potentially damage surface structures, buried installations and plant roots. warming of the ground may cause vapour migration and progressive drying and perhaps even shrinkage of soils. significant heating of the ground can also lead to consolidation and settling in clayey, unconsolidated soils and thermal interference between neighbouring ground-source energy schemes may occur. in the danish legislation for closed-loop systems safety distances to other ground heat exchangers and to extraction wells for drinking water are specified, but not to buildings and other structures. for openloop systems the limitations in the allowable thermal influence on the groundwater are rather strict. evaluation of thermal properties of danish sediments in denmark vertical systems are normally drilled to depths of 50–200 m. in order to evaluate the possible energy extraction from a specific new site and estimate the required depth and number of boreholes, information about the geology and hydrogeology is crucial. the greater part of denmark is underfig. 2. principles of horizontal (a) and vertical (b) closed-loop, groundsource heat-pump systems. 100 km miocene oligocene eocene paleocene cretaceous chalk older than chalk paleocene limestone a 55°n 57° 10°e 14°e 100 km 91–100 81–90 71–80 61–70 51–60 41–50 31–40 21–30 11–20 1–10 0 176–320 151–175 126–150 101–125 thickness (m)b fig. 3. maps showing the pre-quaternary geology (a) and the thickness of quaternary sediments (b) in denmark. ba 39 lain by a sedimentary basin dominated by shallow marine to deep marine clastic and biogenic sediments (fig. 3a), overlain by quaternary deposits. in some areas the quaternary cover is thin and limestone or tertiary sand, silt and mud are found close to the surface. thick successions of quaternary sediments are common in the northern and south-western part of the country (fig. 3b). in these areas the quaternary sediments are dominated by marine and glaciomarine deposits. thick successions of quaternary deposits are also found in buried valleys (jørgensen & sandersen 2006). the geological variation within the quaternary sediments is high and regionalscale geological models are not sufficiently detailed to estimate the possible energy extraction from a vertical system. local information from boreholes and geophysical data are necessary to map the quaternary strata and the elevation of the water table. the potential flow of energy in and out of a vertical system is determined by the heat capacity and the thermal conductivity of the sediments or rocks surrounding the borehole. the process is dominated by heat conduction, but advective transport in the groundwater also plays a role. the thermal conductivity of geological materials depends on the porosity, water content and mineral composition. rocks and sediments with a high content of quartz have high thermal conductivity, whereas porous, unsaturated sediments have low thermal conductivity due to the presence of air in the pore space. thermal conductivity values for different rock and sediment types are compiled in table 1 and show large variations for saturated clay and silt as well as for till deposits, which are common sediments in denmark. most of the values are from germany and the uk and to get more information about common danish lithologies new measurements are required. modelling of near-surface heat flow in a danish context various model systems are available for calculating the efficiency of vertical systems in terms of energy extraction, ranging from simple and easy-to-use models to complex numerical models. the easy-to-use models consider the longterm aspects assuming homogeneous and constant subsurface thermal properties, whereas the complex models also take into account the effect of alternating system operation, heterogeneity in thermal properties as well as groundwater flow. assuming standard design and 1800 hours of production per year, specific heat extraction rates for selected sediment and rock types have been estimated for german (vdi 2001) and british (mcs 2011) conditions based on vertical systems with doubleand single-loop pipes, respectively (table 1). the values indicate large variations between different sediment types and within specific geological facies as well as between the german and british studies, stressing the importance of establishing valid standard values applicable for danish conditions and different systems. the possible energy extraction for 100 m vertical, closed-loop boreholes has been calculated for four simple geological successions (fig.  4). the estimated energy extraction for the most unfavourable of the geological scenarios (7560 kwh) is only c. 60% of the energy extraction for the most favourable (12 537 kwh), thus giving an indication of the possible effect of different geological settings. in addition to heterogeneity in geology and thermal properties, the efficiency of vertical systems is affected by possible groundwater flow. if the system is operating only in heating or cooling mode, groundwater flow will have a positive impact on the efficiency due to the continuous supply of energy. on the other hand, groundwater flow will have a negative sediment/rock thermal conductivity recommended values estimated specific heat extraction rate (w m–1) w mk–1 w mk–1 vdi (2001) mcs (2011) * vdi (2010). † balling et al. (1981). ‡ porsvig (1986). § banks (2008). clay and silt (dry) 0.4–1.0* 0.5* – – water-saturated clay and silt 1.1–3.1* 1.8* 35–50 21–34 palaeogene clay, denmark 1.34–1.56† – – – sand (dry) 0.3–0.9* 0.4* 25 – water-saturated sand 2.0–3.0* 2.4* 65–80 26–45 water-saturated gravel 1.6–2.5* 1.8* – – till/loam 1.1–2.9* 2.4* – – clayey till, denmark 2.00–2.31‡ – – – chalk, england 1.79 ± 0.54§ – – – chalk, denmark 1.45–1.86† – – – quartzite 5.5–7.5§ 6.0§ – – granite 3–4§ 3.4§ 65–85 33–45 table 1. thermal conductivity of different sediment and rock types 4040 impact on systems using the subsurface for seasonal energy storage, because the flow will transport heat and cold away from the vertical system. to evaluate the importance of the hydrogeological heterogeneity observed in denmark, modelling will be carried out based on a complex numerical model that allows the inclusion of subsurface heterogeneity as well as groundwater flow. the aim of the modelling is to identify the aspects that must be considered in the design phase of closed loop systems, and estimate the maximum energy that can be extracted under different hydrogeological conditions typical in denmark. perspectives for shallow geothermal energy in denmark shallow geothermal energy is a competitive, renewable energy resource as domestic heat pumps typically produce 3–4 times the amount of energy they consume in the form of electricity. depending on how the electricity is generated and taking the generation and transmission loss into consideration, heating a building using ground-source energy produces only half of the co2 emission of fossil fuels used directly for heating (banks 2008). nevertheless, the application of shallow geothermal energy in denmark is relatively limited. in 2008, the number of ground-source heat-pump installations was c. 25 000 (willumsen 2008), with the vast majority being horizontal systems. so far, only a few hundred of the installed systems are based on closed-loop boreholes, and the number of groundwater-based open-loop systems is limited to a few tens. hopefully the geoenergy project can pave the way for a wider use of shallow geothermal energy, but in order to fully exploit the resource, a systematic assessment of the potential is needed, including the potential for seasonal storage of waste heat. with the continuous focus on climate change and energy consumption, the number of installations is likely to increase rapidly in the coming years, especially after a political decision has been made to phase out domestic oil burners starting in 2012. the manifold exploitation of various shallow geological resources, including shallow geothermal energy, drinking water and minerals, will increase the demand for resource management and data availability. acknowledgement the eudp programme of the danish energy agency is thanked for financial support to the geoenergy project. references balling, n., kristiansen, j.i., breiner, n., poulsen, k.d., rasmussen, r. & saxov, s. 1981: geothermal measurements and subsurface temperature modelling in denmark. geoskrifter 16, 176 pp. balling, n., nielsen, s.b., christiansen, h.s., christensen, l.d. & poulsen, s. 1992: the subsurface thermal regime and temperature of geothermal reservoirs in denmark. synthesis report to the commission of the european communities, contract en3g-0029-dk, 91 pp. århus: department of earth sciences, university of aarhus. banks, d. 2008: an introduction to thermogeology: ground source heating and cooling, 351 pp. oxford: blackwell. ditlefsen, c. 2012: d1 geologi og jordvarmeboringer. oversigt over geologiske forhold af betydning ved etablering af jordvarmeboringer. geoenergi – energianlæg baseret på jordvarmeboringer – udvikling af markedsfremmende værktøjer og best practice. eudp projekt, j.nr. 64011-0003. jørgensen, f. & sandersen, p.b.e. 2006: buried and open tunnel valleys in denmark – erosion beneath multiple ice sheets. quaternary science reviews 25, 1339–1363. mathiesen, a., kristensen, l., bidstrup, t. & nielsen, l.h. 2009: vurdering af det geotermiske potentiale i danmark. danmarks og grønlands geologiske undersøgelse rapport 2009/59, 30 pp. mcs (microgeneration certification scheme) 2011: microgeneration installation standard: mis 3005. issue 3.0. london: department of energy and climate change. porsvig, m. 1986: varmeovergangsforhold omkring jordslanger. energiministeriets varmepumpeforskningsprogram 33, 56 pp. vdi (verein deutscher ingenieure) 2001: thermische nutzung des untergrundes: erdgekoppelte wärmepumpenanlagen. richtlinie 4640, blatt 2, 43 pp. düsseldorf: verein deutscher ingenieure. vdi (verein deutscher ingenieure) 2010: thermische nutzung des untergrundes: grundlagen, genehmigungen, umweltaspekte. richtlinie 4640, blatt 1, 33 pp. düsseldorf: verein deutscher ingenieure. willumsen, b. 2008: jordvarmeanlæg. teknologier og risiko for jordog grundvandsforurening, 176 pp. miljøprojekt nr. 1238. copenhagen: miljøstyrelsen. fig. 4. four simple geological successions with average, specific heat-extraction rates from table 1 vdi (2001). the possible energy extraction has been calculated to 12 537; 10 373; 9576 and 7560 kwh (left to right), based on 1800 hours of production per year (borehole length × specific heat extraction rate × hours of production). authors’ address geological survey of denmark and greenland, lyseng allé 1, dk-8270 højbjerg, denmark. e-mail: tvp@geus.dk d ep th ( m ) 0 50 100 sand, unsaturated, 25 w m–1 sand, saturated, 72 w m–1 clay, saturated, 42 w m–1 geological survey of denmark and greenland bulletin 23, 2011, 45–48 45 chloride (cl) from dissolved salt is a major threat to groundwater quality in many regions of the world. in arid regions near present-day coastlines, where old seawater occurs in deeper sediments and where road salt is frequently used, cl can be a significant pollutant (european environmental agency 2009). european union member states have recently reported that next to nitrogen, cl is the most commonly found pollutant and is often responsible for groundwater bodies being at risk or having a poor ecological status (european commission 2010). intrusion of salty groundwater and infiltration by seawater near coastlines are well-known phenomena in danish aquifers (ødum & christensen 1936; bonnesen et al. 2009). saltwater in aquifers may also come from human pollution such as landfills, road salt storage facilities, roads and agricultural activities (panno et al. 2006). since the 1970s, deicing salt applied to roads has been recognised as a significant source of contamination that may deteriorate aquifers that are used as drinking water resources by increasing their cl concentration and by harming stream and lake ecosystems (jackson & jabbogy 2005). recent studies have suggested that the decade-long usage of road salt is becoming a rising threat to groundwater quality (bester et al. 2006). salt contamination from roads is therefore particularly problematic to aquifers already at risk (lundmark & olofsson 2007). the present paper explores the impact of road salt on groundwater quality in denmark by means of a combination of chemical indicator analysis, temporal and spatial cl analysis and numerical groundwater modelling. the vulnerability of aquifers to road salt depends on the amount of salt applied per kilometre road, the degree of urbanisation and the percentage of salt lost to the subsurface. based on a literature review, the estimated percentage of road salt lost to the groundwater is 10–20% of applied de-icing salt in danish urban areas (kristiansen et al. 2009). the average amount of road salt used during the winter in denmark has risen since the late 1990s, but varies with weather conditions (fig. 1). in comparison, fig. 1 also shows that the atmospheric nacl deposition for the entire danish surface area is about 280 gigagrams (gg) per year, which is tentatively estimated based on actual bulk deposition measurements (t. ellerman, personal communication 2011). thus, the total amount of applied road salt and the total atmospheric salt deposition in denmark are of the same order of magnitude. however, the local surface load of road salt and atmospheric salt deposition vary widely across denmark. the danish road salt project this paper addresses the results from a recent assessment performed under the danish groundwater mapping project in order to evaluate the impact and risk of road salt to the quality of danish groundwater resources (kristiansen et al. 2009). the project used groundwater quality data from the national database jupiter. the data were downloaded in october 2008 and included approximately 140 000 analyses from approximately 24  000 groundwater sampling points corresponding to about one sample per 2 km2. the oldest data are from 1890, the most recent from 2007. data on the historical consumption of road salt were drawn from the danish road directorate and involved municipalities. three different methods were used: (1) evaluation of indicators to separate chloride sources in groundwater chemistry, (2) analysis of the distribution and variation of cl in time and space, and (3) development of a numerical groundwater risk assessment tool in the hydrological modelling system mike she. does road salt affect groundwater in denmark? søren m. kristiansen, flemming d. christensen and birgitte hansen 500 400 300 200 100 0 1970 year 200019901980 sa le o f ro ad s al t p er w in te r se as o n total sale of road salt in denmark nacl deposition (1 0 0 0 m et ri c to n n es /y ea r) fig. 1. sale of road salt for highway use in denmark from 1965/1966 to 2008/2009 in 1000 metric tonnes (gg) per winter season based on data from the danish road directorate. atmospheric nacl deposition in gg per year is a tentative estimate for the entire danish surface area based on bulk deposition measurements by t. ellerman, university of aarhus, 2011. © geus, 2011. geological survey of denmark and greenland bulletin 23, 45–48. open access: www.geus.dk/publications/bull 4646 chloride source indicators geochemical tools with mass ratios of especially the halides have proved useful for identifying different groundwater cl sources (davies et al. 1998). groundwater salt origins were identified by a graphical technique that distinguishes between multiple sources (panno et al. 2006). this approach discriminates sea salt from, for instance, vacuum salt used as road salt. an average of 40% of danish de-icing salt is vacuum salt. both cl and br form stable anions in water, which are usually not affected by sediment–water reactions. in addition, nabr is less soluble than nacl. as a consequence of the production process, vacuum salt has a cl/br mass ratio >1000. sedimentary rock salt, and hence road salt coming from this source, has cl/br mass ratios similar to rain, as well as residual and infiltrating waters with a cl/br mass ratio <400 (davis et al. 1998). based on a literature review and own data analyses, we find that cl/br ratios are appropriate to detect the origin of dissolved cl sources in danish groundwater (kristiansen et al. 2009). the chemical indicator analysis shows that the potential impact of road salt on groundwater can be traced using a combination of cl/br mass and na/cl molecular ratios in groundwater where the na/cl ratio is affected by the exchange of sodium between the solid and liquid phases and the cl/br ratio is affected by the above-mentioned dissolution processes. groundwater chloride sources firstly, a prevalence of br-poor groundwater (cl/br mass ratios >1000) was found in upper groundwater (<80 m below surface), which indicates that anthropogenic cl sources (e.g. vacuum salt from roads, atmospheric deposition, or animal manure from farming) have a general impact on groundwater quality. secondly, groundwater with reversed ionic exchange (na/cl molecular ratio <0.75) was also preferentially found in upper groundwater, which indicates infiltration of nacl containing water into more fresh sediment. thirdly, most chloride in groundwater (mg/l) >600 250–600 125–250 75–125 30–75 <30 50 km 10°e 14°e 55°n 57°n 57°n 55°n fig. 3 fig. 2. geographical distribution of the latest analysed chloride concentration in c. 24 000 danish groundwater sampling points. the highest measured chloride concentration in wells with more than one measuring point is shown. data were downloaded from the national database jupiter in october 2008. 47 of the groundwater with high cl concentrations was found in the upper groundwater with a gradual decrease from the surface to about 80 m below surface. these three different analyses support that the upper groundwater is affected by cl sources at the soil surface. however, the analyses could not identify which specific type of cl source at the surface influences groundwater quality. in addition, the analyses showed that the deeper groundwater (>90 m below surface) often had cl/br mass ratios <550; moreover, a gradual increase in cl concentrations with depth indicated that the primary cl source in deeper danish aquifers should be found in the underlying groundwater with much higher cl concentrations. chloride distribution in danish groundwater the cl concentration classes used in fig. 2 are based on statistical analysis of the distribution of all the cl analyses from danish groundwater where four geochemi cal populations are found: <10, 10–30, 30–600 and >600 mg cl/l. background concentrations of cl in danish ground water is below 30 mg/l. groundwater with cl concentrations above the drinking water standard of 250 mg/l is commonly found close to the coastline, especially in the eastern parts of denmark (fig. 2). however, elevated concentrations of cl are also found in inland aquifers. trend analyses of the cl concentration in the groundwater in the greater copenhagen area show that 38% of the wells have experienced significantly increasing concentrations whereas only 9% have seen significantly decreasing concentrations (95% confidence interval) since the 1960s. median cl concentrations in groundwater rose from 40–80 mg/l in 1965–1978 to 80–160 mg/l in 1994–2007. a predominance of inversed ionic exchanged groundwater indicates that infiltration of salt water into a fresher aquifer comes from anthropogenic influenced sources such as road salt or intrusion by sea water due to drinking water abstraction. numerical modelling of road salt impact a numerical assessment tool was developed for a large part of the greater copenhagen area (274 km2) in order to evaluate the impact of road salt on groundwater quality. we modelled losses of road salt to the environment at catchment scale from 1967 to 2060 (fig. 3). the tool combines a surface load model with a well-calibrated 3d numerical groundwater model in mike she that simulates water flow and solute transport in the subsurface (kristiansen et al. 2009). the surface load model consists of (1) the historic use of road salt since 2001 distributed on the road network where the type of road has been taken into consideration, (2) estimation of the loss of road salt to the groundwater, and (3) simple 1d modelling of the cl transport through the unsaturated zone. the loss of road salt to the surroundings is difficult to estimate, but tvedt et al. (2001) estimated that 15–30% of the road salt is lost under danish conditions. however, not all the lost road salt infiltrates the groundwater as some percolating water is removed by drainage or sewage. in the modelling, we decided to use a loss of 15% of the road salt to groundwater, which can be considered as a best estimate based on available knowledge. simulations indicate that with a loss of 15% of the applied road salt, the chloride concentrations below urban areas genc h lo ri d e co n ce n tr at io n (m g/ l) well 250 100 25 12.5 3 km rødovre glostrup albertslund brøndby copenhagen frederiksberg copenhagen vallensbæk abstraction well (fig. 4) model boundary hvidovre nn fig. 3. modelling results of chloride concentrations in the primary aquifer due to leaching of road salt estimated for a steady-state situation in 2060 in the greater copenhagen area. the loss of historically used road salt is put to 15% and all factors are kept constant from 2008 to 2060. the background concentration of chloride is not included. for location see fig. 2. 4848 erally will show a 25–40 mg/l increase, whereas increases can reach 125 mg/l at some major road junctions. simulated breakthrough curves for a shallow well are shown in fig. 4 for different scenarios compared to measured values. a background cl concentration of 80 mg/l is added for the simulated results which contains cl from natural sources as atmospheric deposition and marine residual water. the upper curve, where the current load of road salt is maintained, shows that it takes decades before a steady state situation is reached. the rest of the curves show the development of the groundwater cl concentrations at steady state for different reduction scenarios in relation to current road salt usage. conclusions the results show that the upper groundwater cl concentrations (<80 m below surface) are affected by cl sources such as road salt, atmospheric deposition and animal manure. precise identification of the cl sources at the surface requires more analyses of chemical indicator species in the groundwater. numerical groundwater modelling in the greater copenhagen area shows that road salt can result in a significant increase of the cl concentration in groundwater, particularly near major road junctions. the aquifer used for water supply may be degraded because of the accumulated impact from several cl sources such as road salt, residual salt groundwater and recent seawater intrusion. the applied model assumes that the loss to groundwater of road salt is 15%. but if the loss was 30%, then the resultant cl concentration should be doubled. more precise quantification of the loss of road salt and knowledge on cl sources other than salt applied to public roads are therefore required in order to reduce the uncertainty of the current estimate of the effect of road salt on groundwater quality, for example by establishing study sites in urban areas. acknowledgement the project was supported by the former environmental centres under the danish ministry of the environment. references bester, m.l., friend, e.o., molson, j.w., rudolph, d.l. 2006: numerical investigation of road salt impact on an urban well-field. ground water 44, 165–175. bonnesen, e., larsen, f., sonnenborg, t., klitten, k. & stemmerik, l. 2009: deep saltwater in chalk of north-west europe: origin, interface characteristics and development over geological time. hydrogeology journal 17, 1643–1663. davis, s., cecil, d.w., zreda, m. & sharma, p. 1998: uses of chloride/ bromide ratios in studies of potable water. ground water 36, 338–350. european commission 2010: report from the commission in accordance with article 3.7 of the groundwater directive 2006/118/ec on the establishment of groundwater threshold values, 10 pp. brussels: european commission. european environmental agency 2009: water resources across europe – confronting water scarcity and drought, 60 pp. copenhagen: european environmental agency. jackson, r.b. & jabbogy, e.g. 2005: from icy roads to salty streams. proceedings of the national academy of sciences of the united states of america 102, 14487–14488. kristiansen, s.m., christensen, f.d. & hansen, b. 2009: vurdering af danske grundvandsmagasiners sårbarhed overfor vejsalt, 107 pp. københavn: de nationale geologiske undersøgelser for danmark og grønland. lundmark, a. & olofsson, b. 2007: chloride deposition and distribution in soils along a deiced highway – assessment using different methods of measurement. water, air & soil pollution 182, 173–185. ødum, h. & christensen, w. 1936: danske grundvandstyper og deres geologiske optræden. danmarks geologiske undersøgelse iii. række 26, 183 pp. panno, s.v., hackley, k.c., hwang, h.h., greenberg, s.e., krapac, i.g., landsberger, s. & o’kelly, d.j. 2006: characterization and identification of na-cl sources in ground water. ground water 44, 176–187. tvedt, t., randrup, t.b., pedersen, l.b. & gludsted, s. 2001: planter & vejsalt, 19 pp. københavn: trafikministeriet, vejdirektoratet and miljøog energiministeriet, skov & landskab. current average road salt load maintained 25% reduction of current road salt load 50% reduction of current road salt load 130 120 110 100 90 80c h lo ri d e co n ce n tr at io n ( m g/ l) year 1980 2000 2020 2040 2060 100% reduction of current road salt load no road salt used on highways; current road salt load maintained on other road types observed data fig. 4. modelled and measured chloride concentrations in groundwater in an abstraction well in the greater copenhagen area. for location see fig. 3. breakthrough curves are shown for different scenarios whereby road salt losses to groundwater are reduced. authors’ addresses s.m.k., university of aarhus, department of earth sciences, høegh-guldbergs gade 2, dk8000 aarhus c. e-mail: smk@geo.au.dk f.d.c., rambøll, hannemanns allé 53, dk-2300 copenhagen s, denmark. b.h., geological survey of denmark and greeland, lyseng allé 1, dk-8270 højbjerg, denmark. geological survey of denmark and greenland bulletin 13, 2007, 57-60 57 earthquake activity in greenland has been registered and mapped since 1907 (larsen et al. 2006) and thus a long (albeit relatively sparse) record of seismic activity is available for evaluation of seismic hazard and risk. seismic hazard assessment is carried out by judging the probability of future earthquakes in a given region and is based on statistic treatment of earthquake data. the determination of the seismic hazard is the first step in an evaluation of seismic risk, i.e. the possible economic costs and loss of human life after an earthquake. the motivation for this seismic hazard study is the registration of four significant earthquakes in greenland in 2005. the geological survey of denmark and greenland (geus) received reports of all four earthquakes from residents who had felt the shaking. the 2005 earthquakes were located at or near qeqertarsuaq on 30 march, sisimiut on 23 july, station nord on 30 august and attu on 23 october (fig. 1), with magnitudes on the richter scale of 4.3, 4.1, 5.1 and 2.5, respectively. the earthquake in attu led to the inhabitants fleeing in their boats. earthquake activity seismic hazard is just one of many natural hazards in green land. other natural hazards include: continuous permafrost that constitutes a serious obstacle to development over the northern two-thirds of greenland; strong katabatic winds that occur along the edge of the ice cap; very low wind chill and cold sea water, and; landslides and landslide-generated tsunamis (e.g. dahl-jensen et al. 2004). offshore geohazards have been studied by geus for the 2002 licensing round off the greenland west coast (christiansen et al. 2002). a first estimate of the seismic hazard of greenland was presented during the global seismic hazard assessment program (gshap; giardini et al. 1999). for greenland gshap used the very sparse data set that was collected prior to the mid-1990s. since that time the data set has been much improved. the earthquake information used in this seismic hazard study has been extracted from the geus earthquake database and constitutes 227 events that occurred from november 1971 to february 2006 (figs 1, 2). the majority of these events had a magnitude between 3.0 and 5.0, but since 2005 improved analytical methods have lowered the detection threshold from 3.0 to 1.0 in some areas. the earthquakes were primarily shallow; 93% have been located between 0 and 40 km depth. the location of the earthquakes in the geus earthquake database is determined primarily by using measurements from the network of permanent and temporary seismic staseismic hazard assessment of greenland peter voss, stine kildegaard poulsen, sebastian bjerregaard simonsen and søren gregersen © geus, 2007. geological survey of denmark and greenland bulletin 13, 57–60. available at: www.geus.dk/publications/bull fig. 1. map showing the nine seismic source zones chosen for seismic hazard assessment of greenland and the location and magnitude of earth quakes registered in the period november 1971 to february 2006 used in this study. in zones bordered by dashed lines the number of earthquakes is too low to provide input values to hazard assessment (see table 1). tions. by the end of 2006, the network of broadband, seismic stations in greenland included four permanent and 14 temporary stations (larsen et al. 2006). if an earthquake in greenland has been recorded by a seismic network in a neighbouring country, these recordings are included in the database. the neighbouring networks that have provided most data are operated by the geological survey of canada, the norwegian seismic array (norsar) and the university of bergen, norway. data analysis the input for this seismic hazard study is the location, time and magnitude of the observed earthquakes. the location and time are determined using a 1d earth model and an iterative linear inversion scheme (lienert & havskov 1995; havskov & ottemöller 2003). the magnitude used is the body wave magnitude mb, and if that is not available, the local magnitude ml (gregersen 1982). the attenuation of seismic lg waves is known at a few seismic stations as de scribed by gregersen (1982), but a full attenuation model for greenland is at present not available. we have therefore applied the global reference model sea96 by spudich et al. (1997) that describes attenuation from normal faults in hardrock conditions. focal plane solutions have been estimated for only five earthquakes in greenland (e.g. gregersen 2006), and have thus not been included in this study. palaeoseismic information could be of interest if available, but has not been included since palaeoseismic data are of little influence for return periods less than 1000 years in intraplate settings (ata kan et al. 2001). offshore reflection seismic profiles along the west coast of greenland (chalmers & pulvertaft 2001) show the presence of major fault systems that should be taken into account for return periods longer than 1000 years. the coastal area of greenland has been divided into eight seismic source zones that were chosen according to geological structures and the seismicity of the area (fig. 1). the central part of the ice cap is represented by a single zone. input values for the hazard assessment for each seismic source zone were determined from b-value and magnitude estimates (see table 1 and fig. 3). as an example, the computation of the b-value for seismic source zone 4 is shown in fig. 3. the hazard computation is in the form of estimated maximum acceleration for a return period of 475 years, and is described in detail by poulsen & simonsen (2006). ice cap seismic hazards earthquake measurements on the greenland ice cap have shown that the seismic energy is transmitted through the ice for both local and teleseismic earthquakes. the attenuation model we have applied to estimate the seismic hazard in green land is based on hard-rock conditions; thus, the approach used in this study and the results obtained do not apply for conditions on the ice cap. ground shaking from the newly discovered glacial earthquakes, or icequakes, (ekström et al. 2003, larsen et al. 2006) has not been taken into account either. only a very few earthquakes have been located below the ice cap (see fig. 1), and seismic hazard in the interior of greenland is therefore considered to be even lower than the lowest hazard in any coastal area. 58 fig. 2. seismogram for the richter scale 2.5 attu earthquake on 23 october, 2005 recorded at the seismic station in kangerlussuaq. each trace shows the ground velocity in the direction indicated. the seismogram is band-pass filtered between 1 and 5 hz, and the scale of the maximum amplitude is given at the end of each trace. seismic hazard the seismic hazard assessment of greenland is computed for a return period of 475 years as shown in fig. 4. the maximum hazard is found in seismic source zone 4 at a value of 0.051 g (50.37 cm/s2). from this result we assess the general seismic hazard in greenland to be low, following the classification of jiménez et al. (2003) in which low, moderate and high hazards correspond to peak ground accelerations of 0.0–0.08 g, 0.08–0.24 g and above 0.24 g, respectively, for a 475-year return period. seismic source zone 4 covering the northern and northeastern parts of greenland is the area with the highest seismic hazard; the seismic hazard is below 0.05 g in the other seismic source zones where the highest hazards are encountered in the disko bugt – sisimiut area (seismic source zone 8) followed by southern greenland (seismic source zone 1). these results differ considerably from the gshap estimates that are below 0.02 g for the northern and the north-eastern parts of greenland. along the east coast of the baffin bay we find the seismic hazard to be below 0.024 g, where gshap reported a hazard of up to 0.08 g. seismic risk a full evaluation of seismic risk would include collecting indepth knowledge of factors such as infrastructure, building standards and population density distribution, work that is beyond the scope of this study. here we outline only the expected, relative seismic risk for the four areas of highest seismic hazard. though the highest seismic hazard is found in the northern and north-eastern parts of greenland (seismic zone 4), the seismic risk is very low, since the only permanent residents are the five danish air force personnel at station 59 fig. 3. histogram showing the number and magnitude of earthquakes determined in seismic source zone 4. it is a global experience in seismology that the number of earthquakes (n) of various magnitudes (m) can be expressed in a logarithmic relation: log n = a – b × m (black line). the a-value is a measure of regional seismicity, whereas the b-value may be viewed as a regional physical constant. fig. 4. map showing seismic hazard in greenland for a 475-year return period, corresponding to the 10% probability of exceeding a given gvalue in a 50-year period (hard-rock conditions). seismic source zones from fig. 1 indicated. 60 nord and the staff at the danmarkshavn weather station, where the buildings are designed to withstand the arctic climate. the next two zones in order of decreasing seismic hazard are the disko bugt – sisimiut area (seismic zone 8) and southern greenland (seismic zone 1). larger infrastructure and denser population in seismic zones 1 and 8 indicate that these are judged to be at the highest seismic risk in the whole of greenland. concluding remarks the results presented in this study have completely changed the seismic hazard assessment of greenland compared to the seismic hazard map compiled by gshap (giardini et al. 1999). gshap overestimated the seismic hazard of the area north of disko bugt and underestimated the seismic hazard for the southern, the northern and the north-eastern parts of greenland. during recent years, the majority of reports received by geus of earthquakes felt by the resident population are from tasiilaq (seismic zone 2), but our results show that the seismic hazard in this area is low because these earthquakes are small. the overall seismic hazard in greenland is low compared to the many other natural hazards. however, large destructive earthquakes can occur unexpectedly even in areas with low seismicity, as illustrated by the magnitude 6.3 latur earthquake in central india on 30 september, 1993 (gupta 1993). acknowledgements geoforschungszentrum (gfz) potsdam, germany, provides instrumentation and technical support to the seismograph in danmarkshavn and, together with the incorporated research institutions for seismology (iris), usa, to the seismograph in kangerlussuaq. gfz has also contri buted to the installation and operation of several temporary seismic stations. the bureau of minerals and petroleum, government of greenland, provided financial support for several temporary installations. references atakan, k., ojeda, a., camelbeeck, t., & meghraoui, m. 2001: seismic hazard analysis results for the lower rhine graben and the importance of paleoseismic data. netherlands journal of geosciences 80, 305–314. chalmers, j.a. & pulvertaft, t.c.r. 2001: development of the continental margins of the labrador sea: a review. geological society (london) special publications 187, 77–105. christiansen, f.g., bojesen-koefoed, j.a., chalmers, j.a., dalhoff, f., marcussen, c., nielsen, t., nøhr-hansen, h. & sønderholm, m. 2002: petroleum geological activities in west greenland in 2001. geology of greenland survey bulletin 191, 57–60. dahl-jensen, t., larsen, l.m., pedersen, s.a.s., pedersen, j., jepsen, h.f., pedersen, g., nielsen, t., pedersen, a.k., von platen-hallermund, f. & weng, w. 2004: landslide and tsunami 21 november 2000 in paatuut, west greenland. natural hazards 31, 277–287. ekström, g., nettles, m., & abers, g.a. 2003: glacial earthquakes. science 302, 622–624. giardini, d., grünthal, g., shedlock, k.m., & zhang, p. 1999: the gshap global seismic hazard map. annali di geofisica 42, 1225–1230. gregersen, s. 1982: seismicity and observations of lg wave attenuation in greenland. tectonophysics 89, 77–93. gregersen, s. 2006: intraplate earthquakes in scandinavia and green land. neotectonics or postglacial uplift. journal of indian geophysical union 10, 25–30. gupta, h.k. 1993. the deadly latur earthquake. science 10, 1666–1667. havskov, j., & ottemöller, l. 2003: seisan: the earthquake analysis software for windows, solaris and linux, version 8.0, 244 pp. bergen: institute of solid earth physics, university of bergen, norway. jiménez, m.-j., giardini, d. & grünthal, g. 2003: the esc-sesame unified hazard model for the european-mediterranean region, emsc/csem newsletter 19, 2–4. larsen, t.b., dahl-jensen, t., voss, p., jørgensen, t.m., gregersen, s. & ras mussen, h.p. 2006: earthquake seismology in greenland – im proved data with multiple applications. geological survey of den mark and greenland bulletin 10, 57–60. lienert, b.r.e. & havskov, j. 1995: a computer program for locating earthquakes both locally and globally. seismological research letters 66, 26–36. poulsen, s.k. & simonsen, s.b. 2006: seismic hazard analysis of greenland and a distribution of earthquakes, 68 pp. unpublished b.sc. thesis, university of copenhagen, denmark. spudich, p. et al. 1997: sea96 – a new predictive relation for earthquake ground motions in extensional tectonic regimes. seismological research letters 68, 71–79. authors’ addresses p.v & s.g., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: pv@geus.dk s.k.p. & s.b.s., niels bohr institute, university of copenhagen, juliane maries vej 30, dk-2100 copenhagen ø, denmark. geological survey of denmark and greenland bulletin 26, 2012, 77-80 77 testing of an automatic earthquake detection method on data from station nord, greenland nasim karamzadeh, peter h. voss and gholam d. javan earthquakes are continuously monitored by a global network of several thousand seismic stations equipped with highly sensitive digital seismometers. the geological survey of denmark and greenland (geus) takes part in it by operating five seismic stations in denmark and 18 in greenland, some of the latter in collaboration with international partners. there are two main ways of detecting earthquakes from digital recordings of seismometers: (1) by a manual review of the data by an expert in processing seismic earthquake signals and (2) by an automatic method that uses a computerised algorithm to analyse the recordings. since the beginning of earthquake recording, earthquake detection at geus has been based on the manual review method. there are several reasons why an automatic detection procedure has not yet been implemented at geus: (1) historically, the staff at geus have conducted high-quality manual detection of earthquakes, based on a long tradition of manual seismogram analysis (lehmann 1954), (2) the ambient noise level in denmark is generally too high for small local earthquakes to be detected automatically and (3) in greenland, the distance between the seismometers is too long for automatic methods. previous tests on geus data showed that automatic detection using the so-called standard method resulted in a very high number of false detections, and the effort needed to distinguish real earthquake signals from noise signals was much greater than that needed in the manual method. in addition, the automatic method detected fewer earthquakes than the manual analysis. therefore, new automatic methods are needed to extract real earthquake signals from the background noise. in this article, we present results from testing a newly developed automatic detection method based on analysis of the frequency content of seismic signals. the aim of the study was to investigate whether the automated method can be used in greenland or whether the manual procedure is still superior. the new method was tested on seismic data from station nord, which was selected because it is located in a region with many earthquakes, and because there are no nearby seismograph stations to support the measurements. the closest station is located at danmarkshavn c. 540 km to the south. therefore, the majority of earthquakes that occur in this region are only registered at station nord. earthquakes in the station nord region station nord is located in eastern north greenland, in a region where a major tectonic factor is the spreading that occurs along the rift zones in the northern north atlantic and the arctic ocean (døssing et al. 2010), and which gives rise to high seismic activity (fig. 1). another tectonic factor is postglacial isostatic rebound that was the source of three major earthquakes in 1971, 1987 and 1993 (chung 2002), with magnitudes of 5.1, 5.5 and 5.4 on the richter scale. apart from the seismicity observed at the rift zone, most of the earthquakes in the region occur to the south and west of station nord (fig. 1; gregersen 1982). the earthquake hazard in the region is low, but represents the highest in greenland (voss et al. 2007). the hazard was illustrated by a strong tremor felt at station nord on 30 august 2005, caused by a magnitude 4.2 earthquake with an epicentre only 20 km away. automatic earthquake detection the standard automatic earthquake detection method is known as the short term average versus long term average (sta/lta) method (havskov & alguacil 2010). the basis of this method is two running time windows that both compute the average amplitude of the signal, one with a short 40°w 8°w16°w 24°w 80°n 84°n 0° 40°w 8°w 8°e station nord rift zone arctic ocean atlantic ocean 3 2 1 magnitude (richter scale) 200 km fig. 1. map of the region between north-eastern greenland and svalbard. red dots: epicentres located by geus using manual detection methods. blue dots: epicentres located using automatic detection method. green dots: located by both methods. © 2012 geus. geological survey of denmark and greenland bulletin 26, 77–80. open access: www.geus.dk/publications/bull 7878 duration (sta) and the other with a long duration (lta). the lta represents the stable background noise level of the seismic signal, whereas the sta will increase rapidly at the onset of a seismic signal. the signal is often band pass filtered to lower the ambient noise. the earthquake is detected when the sta/lta exceeds a predefined threshold level (e.g. 10). the basic parameters required for this method are (1) the band pass filter, (2) the duration of sta and lta and (3) the sta/lta threshold level. the disadvantage of this method is that it triggers with all ground shaking that exceeds the threshold level. so, if the sta/lta threshold level is set low to detect small earthquakes, many detections that are nothing but noise or bogus events lead to much additional manual post-processing. a high number of noise detections can be avoided by setting the sta/lta at a higher threshold level, but then important earthquakes may not be detected. in many seismic networks, automatic detection is operated with a low sta/lta threshold level, but to avoid a lot of manual post-processing of detected signals, an additional criterion is applied before they are forwarded to manual processing. it states that the detection is only valid if the signal is also detected by a number of other seismic stations (e.g. 3) within a predefined time window; otherwise the detection is disregarded as being seismic noise. however, in order to use this criterion in the detection of small earthquakes, the network must include a large number of seismic stations located less than 100 km from each other. this is not the case in greenland where the distance between the stations is around 300 km in western greenland and around 600 km in eastern greenland. alternative methods are therefore required, and we have chosen an automatic method based on analysis of the short-time fourier transform of the data. in addition to event detection, the automatic method also provides phase picks, magnitude measurements and azimuth estimates, but the primary aim of this study was to test the performance of the automatic event detections. fig. 2. a: vertical component of a continuous seismogram, which included a signal from an earthquake and several short-duration, high-energy perturbations. b: normalised spectrogram produced by short-time fourier transformation for frequencies of 2–16 hz with 2 hz steps. the colours (scaled to use full colour spectrum) represent an estimate of the energy contained at the given frequency, within a short window (about 0.6 s). at 800–900 s, when an earthquake signal appears in the seismogram, the colours on the spectrogram change, which indicates an order of magnitude increase in the seismic energy over all frequencies. the detection of changes in energy over a proper range of frequencies, corresponding to the frequency content of the earthquake signals, led to the detection of an earthquake. other increases in the energy level are seen at several other times, for example at 600–700 s. these peaks are only seen at low frequencies. c: another representation of the same spectrogram showing the variation of local spectral energy for each frequency band over time. a sharp increase is seen at the onset of the earthquake signal, whereas the perturbation of energy in the background noise has disappeared at higher frequency bands. a b c vertical component time (sec.) time (sec.) 0 100 200 300 400 500 600 700 800 900 1000 1100 1200 1300 1400 1500 1600 16000 2 4 6 8 10 12 14 16 0 200 400 600 800 1000 1200 1400 1600 time (sec.) 0 500 0 50 100 0 100 0 100 0 200 0 0 250 500 500f = 4 hz f = 6 hz f = 8 hz f = 10 hz f = 12 hz f = 14 hz f = 16 hz f = 2 hz 0 10 20 30 40 50 60 –100 –50 0 a m p. ( co u n ts ) 50 90 0 100 200 300 400 500 600 700 800 900 1000 1100 1200 1300 1400 1500 fr eq u en cy ( h z) 79 a spectrogram produced by short-time fourier transform is a very useful tool in seismology because it can provide an image indicating the time at which a burst of energy occurs on a seismogram, in addition to the spectral composition of the signal (gibbons et al. 2008). the event detection algorithm used in this study inspects the temporal variation of the signal spectrogram calculated in frequency bands corresponding to the frequency content of local and regional earthquakes (e.g. 2–16 hz). for detected events, the pand s-phases are picked. an example of a recorded seismogram with an earthquake signal and corresponding spectrogram is shown in fig. 2 where an earthquake is observed on a seismogram at an approximate time of 800 s. obvious changes in the colour of the spectrogram take place over a wide range of frequencies along the time axis, which indicate the arrivals of earthquake energy (fig. 2b). accordingly, the detection of a change in energy pattern over a pre-defined range of frequencies, corresponding to the frequency content of the earthquake signals, leads to the detection of an earthquake. the plots presented in fig. 2c show the variation of energy for each frequency band, corresponding to the above spectrogram. these plots provide another representation of the spectrogram. the problem of detecting an earthquake on seismograms is now reduced to detecting sharp increases in the individual time series representing spectral energy versus time (the plot shown in fig. 2c). to avoid false detections due to seismic noise with a frequency content overlapping the analysed frequencies, only detections made in most of the frequency bands are accepted. for instance, detections should be made at about the same time in at least five out of eight frequency sub-bands for a given spectrogram (fig. 2b). three missing detections are allowed, because this may happen for small events and noisy backgrounds, or low signal to noise ratio in some frequency bands. to reduce the false detection rate, all three components (vertical, n–s and e–w) of the seismograms are used in the event detection procedure. results to test the automatic detection method, station nord data from 6 july 2010 to 6 march 2011 were used. prior to this period, the digitising unit of the seismometer had been upgraded to sample at 100 hz. earlier, the instrument had sampled at 20 hz; this limits earthquake analysis to frequencies below 10 hz, which the automatic method was not prepared for. station nord is equipped with a streckeisen sts-2 sensor and a quanterra q330 digitiser. the automatic method analyses data from all three components of the sensor, using 24 hour data files. the data are band pass filtered between 0.95 and 20 hz before the detection algorithm is applied. the fig. 3. a: seismogram of a magnitude 2.2 earthquake filtered with a 1–17 hz band pass filter containing the frequencies used for the automatic detection. the epicentre was located 354 km south-south-east of station nord. the automatic p-phase (p) was kept in the review, but the automatic s-phase (s) was repicked moving the epicentre 118 km. automatic s pick is seen on the north–south channel, manual s pick is seen on the east–west channel and the automatic tremor duration is seen on the vertical channel. b: vertical component of an earthquake not detected by the automatic method. top trace: data with the 1–17 hz filter used by the automatic method. bottom trace: data with the 2–9 hz filter used by the manual method. the earthquake had a magnitude of 1.1 and was located 234 km east-south-east of station nord. utc: universal time, coordinated. 04:53 04:54 04:55 04:56 04:57 04:58 time (min.; utc) time (min.; utc) 10:17 10:18 10:19 10:20 10:21 filter: 1–17 hz filter: 2–9 hz filter: 1–17 hz plot start 10 feb. 2011 4:52:15.000 vertical component duration north–south component east–west component s s p a plot start 26 sep. 2010 10:15:59.320 vertical component vertical component p b 8080 detections are stored in the international seisan format. during the test period, 156 detections were recorded with the automatic method. a review of these detections showed that 13 were due to sensor calibrations or sensor noise, and 52 were caused by other noise signals such as man-made noise. the remaining 91 events were identified as earthquakes. to test the performance of the automatic method, the data were also analysed with the rasmussen method, where data are manually scanned using predefined time windows and filters (rasmussen & voss 2011). during this manual scanning, the z-channel was checked in two-hourly windows using a 2–9 hz band pass filter, the same filter that is used to analyse the data from the seismometer at danmarkshavn in the daily processing at geus. the scanning found 229 earthquakes in the study period, approximately two and a half times more earthquakes than found by the automatic method. an example of an earthquake that was not detected by the automatic method is shown in fig. 3b. comparison of the lists of earthquakes detected by the two methods shows that the automatic detections do not give a full subset of the manual detections, as the automatic method detected 26 events that were not found by the manual method. a review of these 26 earthquakes showed that 23 had a low signal to noise ratio in the frequency range of the band pass filter (2–9 hz) used and three earthquakes were missed during the analysis. after the detection procedure, we processed the detected events. in this process, the arrival times of the pand s-phases were picked and the magnitude of the signal determined. the automatic method picks pand s-phases and uses the duration of the signal tremor to estimate the magnitude. we reviewed the 91 earthquakes detected and found that the automatic pick of the pand s-phases should be adjusted for all earthquakes. most of the adjustments were within a few seconds, but for seven of the events the automatic pick errors were several seconds. in many cases, the duration of the automatic signal located the end of the event in the last part of the s wave, but in some instances the end of the duration was in the end of the p wave. a similar approach was used by agius & galea (2011) with good results, but further improvement of the processing is required. examples of the performance of the automatic method are shown on seismograms in fig 3. discussion and conclusions the manual detection method found two and a half times more events than the automatic detection method. replacement of the manual method with the automatic detection of earthquakes in greenland will therefore result in a significantly lower number of earthquake detections. the quality of earthquake risk evaluation depends on a high detection level; thus a lower number of detections will lead to a lower quality of an earthquake risk evaluation. the automatic method contributed with 26 (11%) additional events to the manual detections, and hence a combination of the two methods may increase the number of earthquakes detected in the region. in addition, the automatic method was effective in avoiding false detections. the automatic method gives earthquake locations and magnitude estimates, based on automatic phase picks, phase polarisation and signal duration, but the quality of this information is poorer than that obtained by manual processing. a manual review of the data will still be an important part of the quality control. acknowledgements this study is part of a phd project funded by the international institute of earthquake engineering and seismology for the first author, who was a visiting phd student at geus in 2011. the greenland ice sheet monitoring network is thanked for upgrading station nord. references agius, m.r. & galea, p. 2011: a single-station automated earthquake location system at wied dalam station, malta. seismological research letters 82, 545–559. chung, w.-y. 2002: earthquakes along the passive margin of greenland: evidence for postglacial rebound control. pure and applied geophysics 159, 2567–2584. døssing, a., stemmerik, l., dahl-jensen, t. & schlindwein, v. 2010: segmentation of the eastern north greenland oblique-shear margin – regional plate tectonic implications. earth and planetary science letters 292, 239–253. gibbons, s. j., ringdal, f. & kværna, t. 2008: detection and characterization of seismic phases using continuous spectral estimation on incoherent and partially coherent arrays. geophysical journal international 172, 405–421. gregersen, s. 1982: seismicity and observations of lg wave attenuation in greenland. tectonophysics 89, 77–93. havskov, j. & alguacil, g. 2010: instrumentation in earthquake seismology. modern approaches in geophysics 22, 360 pp. lehmann, i. 1954: characteristic earthquake records. geodætisk instituts skrifter iii. række 18, 33 pp. rasmussen, h.p. & voss, p.h. 2011: detection of earthquakes at the geological survey of denmark and greenland – geus. 42nd nordic seminar on detection seismology, reykjavik, 5–7 october 2011. icelandic meteorological office, programme and abstracts, 41 only. voss, p., poulsen, s.k., simonson, s.b. & gregersen, s. 2007: seismic hazard assessment of greenland. geological survey of denmark and greenland bulletin 13, 57–60. authors’ addresses n.k. & g.d.j., international institute of earthquake engineering and seismology, arghavan st. 21, north dibajee, farmanieh tehran 3913/19395 islamic republic of iran. e-mail: n.karamzadeh@iiees.ac.ir p.h.v., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. geological survey of denmark and greenland bulletin 6, 99-112 99geological survey of denmark and greenland bulletin 5, 99–112 © geus, 2004 jurassic dinoflagellate cyst stratigraphy of store koldewey, north-east greenland stefan piasecki, john h. callomon and lars stemmerik the jurassic of store koldewey comprises a middle jurassic succession towards the south and an upper jurassic succession towards the north. both successions onlap crystalline basement and coarse sediments dominate. three main lithostratigraphical units are recognised: the pelion formation, including the spath plateau member, the payer dal formation and the bernbjerg formation. rich marine macrofaunas include boreal ammonites and the successions are dated as late bathonian – early callovian and late oxfordian – early kimmeridgian on the basis of new collections combined with material in earlier collections. fine-grained horizons and units have been analysed for dinoflagellate cysts and the stratigraphy of the diverse and well-preserved flora has been integrated with the boreal ammonite stratigraphy. the dinoflagellate floras correlate with contemporaneous floras from milne land, jameson land and hold with hope farther to the south in east greenland, and with peary land in north greenland and svalbard towards the north. the middle jurassic flora shows local variations in east greenland whereas the upper jurassic flora gradually changes northwards in east greenland. a boreal flora occurs in peary land and svalbard. the characteristic and stratigraphically important species perisseiasphaeridium pannosum and oligosphaeridium patulum have their northernmost occurrence on store koldewey, whereas taeniophora iunctispina and adnatosphaeridium sp. extend as far north as peary land. assemblages of dinoflagellate cysts are used to characterise significant regional flooding events and extensive sequence stratigraphic units. keywords: ammonites, boreal, dinoflagellate cysts, jurassic, north-east greenland, store koldewey s.p. & l.s., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: sp@geus.dk j.h.c., department of chemistry, university college london, 20 gordon street, london wc1h 0aj, uk. jurassic sediments are well exposed in the east greenland basin from milne land (70°n) in the south to hochstetter forland (75°n) in the north. yet farther to the north, jurassic sediments are restricted to isolated outliers (ravn 1911; stemmerik & piasecki 1990) until the more extensive outcrops of jurassic deposits in the wandel sea basin in eastern north greenland are reached (81°n–83°n; fig. 1). the geographical gap between the jurassic outcrops of the wandel sea basin (peary land) and the well-studied ones of the east greenland basin corresponds to the zone covering the transition from the strictly boreal dinoflagellate cyst flora of peary land to the subboreal hybrid flora of east greenland. accordingly, dinoflagellate data from store koldewey are important since they improve correlation between these two floral provinces. the jurassic exposures on the island of store koldewey are also stratigraphically important because they provide the nearest accessible record in guiding the evaluation of the hydrocarbon potential of the extensive offshore shelf areas in the northern region. the mesozoic of store koldewey was first described by ravn (1911) and koch (1929) on the basis of data collected by members of ‘danmarks expeditionen’ in 1906–1908. store koldewey was then not visited by geologists until the summer of 1989, when several fieldgeus bulletin no 5.pmd 29-10-2004, 11:1499 100 18 w 18ºw 20 km 76º30'n 76ºn store koldewey danmarkshavn undifferentiated, mostly glacial deposits cretaceous kap arendt ravn pynt trækpasset sydlige gneisnæs jurassic caledonian crystalline basement fault n kløft i kløft ii greenland 500 km 5 4 3 2 1 6 1: milne land 2: jameson land 3: hold with hope 4: hochstetter forland 5: store koldewey 6: peary land/ wandel sea basin fig. 1. simplified geological map of store koldewey. geus bulletin no 5.pmd 29-10-2004, 11:14100 101 parties from the geological survey of greenland (ggu; since 1995 part of the geological survey of denmark and greenland) studied the geology of the island as part of regional mapping projects (stemmerik & piasecki 1990; henriksen 1997). other areas mapped as palaeozoic and mesozoic sediments by haller (1983) were also visited at the same time but most of these appeared to be glacial deposits (stemmerik & piasecki 1990). however, some new, very small outcrops of jurassic sediments were found, protected from erosion on the downthrown side of basement faults (piasecki et al. 1994). geological setting the elongated shape of store koldewey reflects a north–south-oriented crystalline basement ridge. mesozoic sediments are exposed only in low coastal cliffs along the east side of the island (fig. 1). they include four distinct stratigraphic units (i–iv) that are preserved in small structural basins separated by basement highs (fig. 1). the lithology is generally of mud to fine-grained sand grade and the units are highly fossiliferous. the stratigraphic units are of middle jurassic (i), late jurassic (ii) and early cretaceous ages (iii–iv). the southernmost outcrop at ravn pynt consists of a middle jurassic succession more than 60 m thick, the trækpasset formation (koch 1929; figs 1, 2). upper jurassic sediments crop out in two gullies, kløft i and kløft ii, in the northern part of the island (fig. 1). the sediments in the first (northern) af these gullies have been referred to the kløft i formation (koch 1929) based on the description by ravn (1911) of material collected by members of the ‘danmarks ekspeditionen’ in 1906–1908. ravn (1911) proposed a callovian age for the trækpasset formation and a ‘sequanian’ (kimmeridgian) age for the kløft i formation based on ammonites. the sandstone of the kløft i formation was later found in the kløft ii locality to be overlain with a sharp boundary by grey laminated mudstones of the bernbjerg formation (piasecki et al. 1994). samples and methods the palynological samples were prepared by standard preparation methods including treatment with hydrochloric and hydrofluoric acids, oxidation and filtration with 20 micron mesh. the stratigraphical position of the middle jurassic samples is marked on the sedimentological log from ravn pynt (fig. 2). this succession is well dated by ammonites and the recorded occurrences and distributions of dinoflagellate cysts are mainly used to improve the precision of middle jurassic dinoflagellate cyst stratigraphy in east greenland. this has involved new palynological analyses of ammonite-dated samples from jameson land for confirmation of the new stratigraphical results from store koldewey. these new data are not published yet, but are referred to in this paper. the new ammonite records are not yet published in detail but are utilised in this paper. the position of the upper jurassic samples is marked on the sedimentological log from kløft ii (see fig. 5). the record of dinoflagellate cysts is then used to date the regionally recorded transition from shallow marine sandstone to the shale of the bernbjerg formation. the dinoflagellate assemblages are correlated with older ammonite data, but the samples were not themselves directly associated with ammonites. a number of new species have been described and defined separately (piasecki 2001). middle jurassic lithostratigraphy the middle jurassic succession on store koldewey, the trækpasset formation of koch (1929), is a lateral equivalent in part of the geographically widespread pelion formation, and as there is little reason to maintain a separate stratigraphy for store koldewey, it is herewith re-assigned to the pelion formation. the upper part of the exposed succession is of early to middle callovian age and biostratigraphically equivalent to the new spath plateau member on hold with hope (vosgerau et al. 2004, this volume) where the same two ammonite zones are recorded. this part of the store koldewey succession provides dinoflagellate cyst data from an interval that is not so well documented in jameson land (milner & piasecki 1996). ammonites and dinoflagellate cysts in the succession at ravn pynt may indicate a depositional break between the bathonian and the callovian parts of the pelion formation (figs 3, 4). the basal callovian comprises the only significant mudstone in this succession. this shift in depositional facies is associated with the appearance of callovian ammonites and several of the uppermost bathonian ammonite zones are not recorded beneath the facies shift. the absence of these zones geus bulletin no 5.pmd 29-10-2004, 11:14101 102 40 0 10 20 30 50 m structure fossils lithology sand sand with gravel mud clay clasts calcareous concretions hardened bed lamination lenticular lamination planar cross-bedding shell bed thalassinoides isp. curvolithos isp. planolites isp. monocraterion isp. diplocraterion isp. boreal ammonite fauna horizon belemnite bivalve pinna sp. serpulid oyster 360691 wood 360691 360694 360695 360696 360697 360699 360392 360393 360394 360395 360396 360397 360398 360399 360400, 401 360700 360698 360692, 693 mud sand gravel 16 16 26 31 31 19a ggu sample numbers fig. 2. sedimentological log of the exposed middle jurassic succession at ravn pynt. levels of the analysed palynological samples are indicated. geus bulletin no 5.pmd 29-10-2004, 11:14102 103 may indicate a hiatus in the sedimentary succession or simply no preservation of fauna and flora in that specific part of the section. some of the absent ammonite faunas, however, have been collected just south of ravn pynt on store koldeway. ammonites the jurassic succession contains abundant ammonites in well-separated faunal horizons and the stratification is generally preserved despite some solifluction. the sedimentological succession was studied in a ravine at ravn pynt and four ammonite horizons were collected here in situ and directly correlated with the fine-grained palynological samples (fig. 2). these ammonites represent the arcticoceras ishmae zone (horizon 16), the arcticoceras cranocephaloide zone (horizon 19a including topotypes of kepplerites tychonis ravn), the cadoceras apertum zone (horizon 26) and the proplanulites koenigi zone (horizon 31) of the standard boreal ammonite biostratigraphy of jameson land (callomon 1993). the age of the succession is then midbathonian tocallovian based on the ammonite faunas. dinoflagellate cysts the organic matter in the middle jurassic succession is dominated by black and brown woody material. the abundance of dinoflagellate cysts is generally low but the diversity is fair (65 recorded species, including a few acritarchs). the assemblages are dominated by proximate cysts accompanied by few cavate cysts but are basically barren of chorate cysts. specimens from the pareodinia/paraevansia/evansia and the escharisphaeridium/sentusidinium groups are the most abundant cysts. successive species appear regularly upwards in the succession but the highest number of appearances is in ggu sample 360698 at 19.60 m (fig. 4) at the base of a 13 m thick fine-grained interval (fig. 2). this may reflect a hiatus in the succession below the sample and/or a shift to a more open marine depositional environment. since the middle jurassic succession is dated in detail by ammonites, the distribution of the dinoflagellate cysts can be used to increase the precision of their ranges in the boreal jurassic as previous recorded in other regions in east greenland (jameson land and hold with hope). sirmiodinium grossii, gonyaulacysta pectinigera, paragonyaulacysta retiphragmata and chytroeisphaeridia hyalina are the stratigraphically important species that occur throughout the investigated succession. s. grossii appears in the arctocephalites arcticus chronozone in milne land (larsen et al. 2003) and becomes more frequent from approximately the arctocephalites cranocephaloide chronozone and upwards to the cadoceras apertum chronozone (milner & piasecki 1996). this increase in abundance of s. grossii is not observed here because no samples were collected from this part of the section. in milne land, g. pectinigera is recorded as becoming abundant from the a. cranocephaloide chronozone upwards into the c. apertum chronozone but again, this increase in abundance is not represented in the present section. in jameson land, p. retiphragmata has its earliest appearance in the a. cranocephaloide chronozone but at store koldewey it clearly appears already in the arcticoceras ishmae chronozone. c. hyalina appeared earliest in the cadoceras calyx chronozone in jameson land (milner & piasecki 1996). the present material, in combination with new data from jameson land, show that c. hyalina is present already in the a. ishmae chronozone and becomes abundant from the c. apertum zone. aldorfia aldorfensis appears in the a. cranocephaloide chronozone precisely where it is expected from its range in jameson land (milner & piasecki 1996). the morphologically variable species ctenidodinium thulium is recorded throughout the succession, having a much more extensive range than that recorded elsewhere in east greenland. towards the south, in milne land, it is restricted to the interval above the a. cranocephaloide chronozone and below the paltoceras athleta chronozone (larsen et al. 2003). this upper range is confirmed in these northern regions by data from hold with hope where it is recorded in strata just below the p. athleta chronozone (piasecki et al. 2004, this volume). many of the less frequent species seem to occur in a rather random way that makes their presence and especially their absence less stratigraphically useful. crussolia perireticulata is a good example of this. in milne land, it occurs in the bathonian a. arcticus chronozone and again higher in the callovian part of the succession (larsen et al. 2003). in jameson land, it appears one ammonite zone higher than in milne land, the bathonian arctocephalites greenlandicus chronozone and is then not recorded again until it reappears in the basal callovian, top c. apertum chronozone. there are no records of this species from the bathonian or the callovian of store koldewey and hold with hope. this could indicate a southern affinity of this geus bulletin no 5.pmd 29-10-2004, 11:14103 104 species as the depositional settings of these localities are comparable overall. however, a callovian range is reported from the present arctic regions of the sverdrup basin and svalbard (smelror 1993). some exceptionally early occurrences of ambonosphaera calloviana and gonyaulacysta helicoidea in the a. ishmae chronozone and of pareodinia stegasta and paraevansia brachythelis in the c. apertum chronozone are recorded in the present succession. the dinoflagellate cysts in the middle jurassic succession are divided into an upper bathonian and a lower callovian dinoflagellate cyst assemblage. the boundary is placed at 19.2 m, above which level a significant number of new species appear (figs 2, 4). most species in the upper bathonian assemblage also occur in the lower callovian assemblage and the composition of the bathonian assemblage varies significantly between the few samples studied. the younger, boreal ammonite zonation lithostratigraphy erymnoceras coronatum succession south of ravn pynt succession at ravn pynt kosmoceras jason sigaloceras calloviense proplanulites koenigi cadoceras nordenskjoeldi spath plateau member pe lio n fo rm at io ncadoceras apertum b at h o n ia n c a ll o v ia n cadoceras calyx cadoceras variabile arcticoceras cranocephaloide arcticoceras ishmae arctocephalites greenlandicus arctocephalites arcticus fa u n a h o r iz o n s 31 22 26 20 19 16 15 hiatus ? fig. 3. schematic stratigraphical classification of the middle jurassic succession on store koldewey. geus bulletin no 5.pmd 29-10-2004, 11:14104 10 5 sa m pl e he ig ht m et re 50 25 0 g g u s am pl e no . 57.00 43.80 39.40 38.60 31.80 29.50 26.50 23.00 19.80 12.40 4.60 3.60 360400 54.50 360401 360397 360395 360394 360393 360392 360700 360699 360698 360696 360694 360693 sy st em ju ra ss ic ( m id dl e) st ag e c al lo vi an ba th on ia n f or m at io n pe lio n fo rm at io n 1 pa ra ev an sia s pp . 2 pa re od in ia s co pa eu s 3 am bo no sp ha er a ca llo vie ns e 4 va le ns ie lla o vu la 5 g on ya ul ac ys ta ju ra ss ic a 6 va le ns ie lla d ic ty di a 7 g on ya ul ac ys ta p ec tin ig er a 8 pa ra go ny au la cy st a re tip hr ag m at a 9 ev an sia ja ne ae 1 0 m ei ou ro go ny au la x sp on gi os a 11 ch yt ro ei sp ha er id ia h ya lin a 12 at op od in iu m h ar om en se 13 rh yn ch od in io ps is cf . c la do ph or a 14 si rm io di ni um g ro ss ii 15 ca dd os ph ae ra h al os a 16 m ei ou ro go ny au la x sp p. 17 ch la m yd op ho re lla c f. ec to ta bu la ta 18 d ur ot rig ia d av ey i 19 g on ya ul ac ys ta h el ic oi de a 20 ba tia ca sp ha er a sp p. 21 sc rin io di ni um s pp . 22 se nt us id in iu m sp p. 23 ch yt ro ei sp ha er id ia c hy tr oe oi de s 24 cy m at io sp ha er a sp p. 25 se nt us id in iu m p el io ne ns e 26 ka llo sp ha er id iu m s pp . 27 ct en id od in iu m th ul iu m 28 pa re od in ia a rc tic a 29 va lva eo di ni um le ne ae 30 es ch ar isp ha er id iu m r ud is 31 d ur ot rig ia c f. da ve yi 32 pa re od in ia p ac hy ce ra s 33 ka llo sp ha er id iu m h yp or na tu m 34 po lys te ph an op ho ru s cf . p ar ac al at us 35 va le ns ie lla s p. f en so m e 36 va le ns ie lla s pp . 37 ch yt ro ei sp ha er id ia s pp . 38 fr om ea to rn al is 39 g on ya ul ac ys ta c f. ce nt ric on na ta 40 ev an sia p er ire tic ul at a 41 pa re od in ia s te ga st a 42 n . p le ga s va r. di ct yo rn at us 43 al do rfi a al do rfe ns is 44 d iss ilio di ni um sp p. 45 pa ra ev an sia b ra ch yt he lis 46 ka llo sp ha er id iu m p ra us sii 47 pa ra go ny au la cy st a sp p. 48 pa re od in ia g ra nu la ta 49 pa re od in ia p ro lo ng at a 50 sc rin io di ni um c f. lu rid um 51 se nt us id in iu m sp . d f en so m e 52 at op od in iu m p ol yg on al is 53 fr om ea s pp . 54 pa re od in ia s pp . 55 pa re od in ia c f. sc op ae us 56 es ch ar isp ha er id iu m s pp . 57 ja ns on ia sp p. 58 so lis ph ae rid iu m a nk yle to n 59 le pt od in iu m sp p. 60 si rm io di ni op sis sp p. 61 g on ya ul ac ys ta s pp . 62 pa re od in ia c f. gr oe nl an di cu m 63 es ch ar isp ha er id iu m s pp . 64 li th od in ia sp p. 65 am bo no sp ha er a cf . c al lo via na 66 g on ya ul ac ys ta c f. he lic oi de a alphabetical species list 43 aldorfia aldorfensis 3 ambonosphaera calloviana 65 ambonosphaera cf. calloviana 12 atopodinium haromense 52 atopodinium polygonalis 20 batiacasphaera spp. 15 caddosphaera halosa 17 chlamydophorella cf. ectotabulata 23 chytroeisphaeridia chytroeoides 11 chytroeisphaeridia hyalina 37 chytroeisphaeridia spp. 27 ctenidodinium thulium 24 cymatiosphaera spp. 44 dissiliodinium spp. 31 durotrigia cf. daveyi 18 durotrigia daveyi 63 escharisphaeridium spp. 30 escharisphaeridium rudis 56 escharisphaeridium spp. 9 evansia janeae 40 evansia perireticulata 53 fromea spp. 38 fromea tornalis 39 gonyaulacysta cf. centriconnata 66 gonyaulacysta cf. helicoidea 19 gonyaulacysta helicoidea 5 gonyaulacysta jurassica 7 gonyaulacysta pectinigera 61 gonyaulacysta spp. 57 jansonia spp. 33 kallosphaeridium hypornatum 46 kallosphaeridium praussii 26 kallosphaeridium spp. 59 leptodinium spp. 64 lithodinia spp. 10 meiourogonyaulax spongiosa 16 meiourogonyaulax spp. 42 n. plegas var. dictyornatus 45 paraevansia brachythelis 1 paraevansia spp. 8 paragonyaulacysta retiphragmata 47 paragonyaulacysta spp. 48 pareodinia granulata 28 pareodinia arctica 62 pareodinia cf. groenlandicum 55 pareodinia cf. scopaeus 32 pareodinia pachyceras 49 pareodinia prolongata 2 pareodinia scopaeus 54 pareodinia spp. 41 pareodinia stegasta 34 polystephanophorus cf. paracalatus 13 rhynchodiniopsis cf. cladophora 50 scriniodinium cf. luridum 21 scriniodinium spp. 25 sentusidinium pelionense 51 sentusidinium sp. d fensome 22 sentusidinium spp. 60 sirmiodiniopsis spp. 14 sirmiodinium grossii 58 solisphaeridium ankyleton 6 valensiella dictydia 4 valensiella ovula 35 valensiella sp. fensome 36 valensiella spp. 29 valvaeodinium leneae store koldewey ravn pynt >50 specimens 20–50 specimens 5–19 specimens 1–4 specimens fig. 4. distribution chart of the dinoflagellate cysts in the middle jurassic succession at ravn pynt on store koldewey. g e u s b ul le tin n o 5. pm d 29 -1 020 04 , 1 1: 14 10 5 106 lower callovian assemblage is characterised by abundantchytroeisphaeridia hyalina and pareodinia pachyceras, together with many species not present below e.g. pareodinia stegasta, aldorfia aldorfiense and paraevansia brachythelis. middle jurassic correlation the dinoflagellate cyst assemblage of the lower part of the succession (bathonian) correlates with assemblages of the a. arcticus – a. cranocephaloide chronozones from both milne land (assemblages 2 and 3 in: larsen et al. 2003) and jameson land (milner & piasecki 1996) in central east greenland. the assemblages from the charcot bugt formation in milne land are very poor, both in diversity and density, but have stratigraphically significant species such as s. grossii, c. hyalina, kallosphaeridium hypornatum and evansia janeae in common with the present assemblage. in the same stratigraphical interval in the fossilbjerget formation of jameson land, new species appear for the first time in abundance and have many species in common with the assemblage from store koldewey. differences in occurrence and first appearances of significant species between the two areas are discussed above. the dinoflagellate cyst assemblage of the higher part of the succession (lower callovian) on store koldewey also correlates well with assemblages from the charcot bugt formation in milne land, whereas the contemporaneous assemblages from the fossilbjerget formation in jameson land are poor and not so well documented. in contrast, dinoflagellate cyst assemblages from the pelion formation, spath plateau member, on hold with hope (piasecki et al. 2004, this volume) correlate well with assemblages from store koldewey. the assemblages in the charcot bugt formation (assemblages 4 and 5 in: larsen et al. 2003) are not precisely dated but are not older than the a. cranocephaloide chronozone (bathonian) and not younger than the erymnoceras coronatum chronozone (top middle callovian). discussion the correlation of the pelion formation on store koldewey with the charcot bugt, the fossilbjerget and the pelion formations in southern parts of the jurassic east greenland basin complex shows that time-equivalent sedimentary successions exist regionally in different lithostratigraphical units. the middle jurassic dinoflagellate assemblages vary with depositional environments but the overall characters can be recognised and correlated over long distances. the main problem is the interdependence of deposition of relatively fine-grained sediments and preservation of dinoflagellate cysts, giving relatively few horizons with rich assemblages in the generally coarsegrained middle jurassic successions, i.e. the pelion and charcot bugt formations. upper jurassic lithostratigraphy the upper jurassic succession on store koldewey was defined as the kløft i formation (koch 1929). its lower sandstone-dominated part is equivalent to the recently defined payer dal formation (alsgaard et al. 2003) from hold with hope, kuhn ø and hochstetter forland further to the south, and the overlying mudstones are equivalent to the geographically extensive bernbjerg formation (fig. 5). for convenience and simplification of the lithostratigraphy, the kløft i formation is not used here and the succession is referred to the payer dal and bernbjerg formations (fig. 5). 0 5 10 15 m 360497 360498 360496 360493 360494 360495 mud sand be rn bj er g fm pa ye r d al fm fig. 5. sedimentological log of the exposed upper jurassic succession at kløft ii with the levels of analysed samples marked. for legend, see fig. 2. geus bulletin no 5.pmd 29-10-2004, 11:14106 107 ammonites the old ammonite collections from the kløft i formation (payer dal formation) described by ravn (1911) were recently correlated more precisely with the british ammonite succession, confirming a late oxfordian to early kimmeridgian age, equivalent to the amoeboceras serratum, a. rosenkrantzi and aulacostephanoides mutabilis zones (fig. 6; sykes & surlyk 1976; sykes & callomon 1979). new ammonite finds indicate a similar age. these ammonites are not precisely located but most probably came from the sandstones referred here to the payer dal formation. the samples analysed for dinoflagellate cysts are from the top of this formation and from the overlying bernbjerg formation (fig. 6). dinoflagellate cysts the organic content of the upper jurassic samples is dominated by brown and black woody material and the abundance and diversity of the dinoflagellate cyst floras are low. the richest samples are from the base of the bernbjerg formation, probably representing a flooding event. the dinoflagellate cysts are better preserved in the payer dal formation than in the bernbjerg formation. the composition of the dinoflagellate cyst flora is clearly in favour of proximate cysts, with a minority of chorate specimens and species. the upper jurassic succession comprises two assemblages, which intermingle at the transition from payer dal to bernbjerg formation (fig. 7). the lower assemblage is characterised by gonyaulacysta dualis, adnatosphaeridium sp. (a. hartzi in: piasecki 1980), taeniophora iunctispina, ambonosphaera calloviana and paragonyaulacysta capillosa. this assemblage is well known from oxfordian/kimmeridgian strata in milne land in east greenland (piasecki 1980; piasecki 1996), hochstetter forland (piasecki & stemmerik 2004, this volume) and in peary land, north greenland (håkansson et al. 1981; piasecki 1994). the upper assemblage is characterised by paragonyaulacysta capillosa, occisucysta sp., perisseiasphaeridium pannosum, rhynchodiniopsis sp. and rhynchodiniopsis cf. pennata. this assemblage is known from kimmeridgian strata in milne land with a slightly different frequency of the species involved (piasecki 1996), and from peary land, north greenland (piasecki 1994) and svalbard (århus 1988). the earliest paragonyaulacysta capillosa on milne land appears in the basal kimmeridgian, rasenia cymodoce chronozone, shortly before the earliest perisseiasphaeridium pannosum and avellodinium cf. falsificum at the base of the a. mutabilis chronozone. these two events are recorded within an assemblage of abundant adnatosphaeridium sp. (a. hartzi in: piasecki 1980), taeniophora iunctispina, gonyaulacysta jurassica/dualis and ambonosphaera calloviana that dominates from the late oxfordian (a. rosenkrantzi based on dinoflagellate cysts based on ammonites chronozones k im m er id gi an o xf or di an lithostratigraphy formations bernbjerg payer dal aulacostephanoides mutabilis rasenia cymodoce pictonia baylei amoeboceras rosenkrantzi amoeboceras regulare amoeboceras serratum amoeboceras glosense chronostratigraphyfig. 6. schematic stratigraphical classification of the upper jurassic succession on store koldewey. geus bulletin no 5.pmd 29-10-2004, 11:14107 10 8 25 0 18.00 11.00 4.00 3.00 1.00 360495 360494 360493 360496 360497 ju ra ss ic k im m er id gi an o xf or di an –k im m er id gi an be rn bj er g pa ye r d al 1 rh yn ch od in io ps is cl ad op ho ra 2 pi lo sid in iu m m yr ia tr ic hu m 3 n um m us s pp . 4 ad na to sp ha er id iu m s p. 5 am bo no sp ha er a ca llo via na 6 pa ra go ny au la cy st a ca pi llo sa 7 le pt od in iu m s ub til e 8 se nt us id in iu m p el io ne ns e 9 g on ya ul ac ys ta c f. he lic oi de a 10 le io sp ha er id ia s pp . 11 ep ip lo sp ha er a sp p. 12 g on ya ul ac ys ta d ua lis 13 cr ib ro pe rid in iu m g ra nu lig er a 14 pa re od in ia h al os a 15 te nu a hy st rix 16 ch yt ro ei sp ha er id ia h ya lin a 17 ep ip lo sp ha er a bi re tic ul at a 18 ci rc ul od in iu m sp p. 19 es ch ar isp ah ae ria la ev ig at a 20 ca dd os ph ae ra c f. ha lo sa 21 o cc isu cy st a af f. m on oh eu ris ko s 22 ta en io ph or a iu nc tis pi na 23 am bo no sp ha er a sp . 24 si rm io di ni um g ro ss ii 25 pa re od in ia s pp . 26 sc rin io di ni um ir re gu la re 27 at op od in iu m h ar om en se 28 pe ris se ia sp ha er id iu m p an no su m 29 av el lo di ni um c f. fa lsi fic um 30 es ch ar isp ha er id ia p oc oc ki i 31 rh yn ch od in io ps is sp p. 32 rh yn ch od in io ps is cf . p en na ta 33 o cc iss uc ys ta s pp . 34 ba rb at ac ys ta p ilo sa 35 pr ol ix os ph ae rid iu m g ra nu lo su m 36 es ch ar isp ah ae ria c f. la ev ig at a 37 tu bo tu be re lla c f. eg em en ii 38 te nu a sp p. 39 ci rc ul od in iu m d ow ni ei 40 ci rc ul od in iu m c f. do w ni ei 41 pa ra go ny au la cy st a cf . c ap illo sa 42 o lig os ph ae rid iu m p at ul um 43 ap te od in iu m s pp . 44 cr ib ro pe rid in iu m c f. pe rfo ra ns 45 pa re od in ia b or ea lis 46 pe ris se ia sp ha er id iu m c f. pa nn os um 47 a cr ita rc h sp ec ie s alphabetical species list 23 ambonosphaera sp. 47 acritarch species 4 adnatosphaeridium sp. 5 ambonosphaera calloviana 43 apteodinium spp. 27 atopodinium haromense 29 avellodinium cf. falsificum 34 barbatacysta pilosa 20 caddosphaera cf. halosa 16 chytroeisphaeridia hyalina 40 circulodinium cf. downiei 39 circulodinium downiei 18 circulodinium spp. 44 cribroperidinium cf. perforans 13 cribroperidinium granuligera 17 epiplosphaera bireticulata 11 epiplosphaera spp. 36 escharispahaeria cf. laevigata 19 escharispahaeria laevigata 30 escharisphaeridia pocockii 9 gonyaulacysta cf. helicoidea 12 gonyaulacysta dualis 10 leiosphaeridia spp. 7 leptodinium subtile 3 nummus spp. 33 occisucysta spp. 21 occisucysta aff. monoheuriskos 42 oligosphaeridium patulum 45 pareodinia borealis 6 paragonyaulacysta capilosa 41 paragonyaulacysta cf. capilosa 14 pareodinia halosa 25 pareodinia spp. 46 perisseiasphaeridium cf. pannosum 28 perisseiasphaeridium pannosum 2 pilosidinium myriatrichum 35 prolixosphaeridium granulosum 32 rhynchodiniopsis cf. pennata 1 rhynchodiniopsis cladophora 31 rhynchodiniopsis spp. 26 scriniodinium irregulare 8 sentusidinium pelionense 24 sirmiodinium grossii 22 taeniophora iunctispina 15 tenua hystrix 38 tenua spp. 37 tubotuberella cf. egemenii sa m pl e he ig ht m et re g g u s am pl e no . sy st em st ag e f or m at io n store koldewey kløft ii >50 specimens 20–50 specimens 5–19 specimens 1–4 specimens fig. 7. distribution chart of the dinoflagellate cysts in the upper jurassic succession at kløft ii on store koldewey. g e u s b ul le tin n o 5. pm d 29 -1 020 04 , 1 1: 14 10 8 109 chronozone) to the earliest kimmeridgian (pictonia baylei and r. cymodoce chronozones). a similar order of appearances occurs in the upper jurassic succession of store koldewey and the presence of occisucysta sp. in these strata also supports a stratigraphical level equivalent to the r. cymodoce chronozone by comparison with the assemblages from milne land. p. pannosum becomes abundant in milne land in the a. mutabilis chronozone and dominates the dinoflagellate assemblages throughout the a. eudoxus zone until oligosphaeridium patulum becomes the totally dominant species in the a. autissiodorensis chronozone. the assemblage on store koldewey contains perisseiasphaeridium pannosum most abundantly at the transition between the two formations and oligosphaeridium patulum follows rapidly at the base of bernbjerg formation. none of the two species are abundant at higher levels in the bernbjerg formation at kløft ii. upper jurassic correlation the upper jurassic dinoflagellate cyst assemblages on store koldewey closely resemble the boreal assemblages from the jurassic ladegårdsåen formation of peary land (piasecki 1994), and to some degree also the assemblage in the janusfjellet formation, svalbard (århus 1988). however, the restricted ammonite faunas in these two formations do not allow precise dating of the dinoflagellate assemblages on peary land and svalbard. the most significant difference between these assemblages is the presence of p. pannosum and o. patulum only at store koldewey. adnatosphaeridium sp. and taeniophora iunctispina occur as far north as peary land but are not reported from svalbard. the stratigraphically highest occurrence of gonyaulacysta dualis, atopodinium haromense, occisucysta sp. in peary land, in an assemblage with abundant escharisphaeridium pocockii, rhynchodiniopsis sp., taeniophora iunctispina and adnatosphaeridium sp. is associated with the appearance of paragonyaulacysta capillosa and cribroperidinium perforans. this assemblage and associated stratigraphic events in peary land are directly comparable with the dinoflagellate cyst assemblage at the transition from the payer dal formation to the bernbjerg formation on store koldewey. the age of this assemblage is interpreted to be early kimmeridgian (r. cymodoce – a. mutabilis chronozones). a comparable event has not been recorded in the kimmeridgian of svalbard (janusfjellet formation) where only paragonyaulacysta capillosa seems to have stratigraphical potential, with a consistently relative short range (århus 1988). there, p. capillosa appears after a poor to barren interval in the upper oxfordian to lowermost kimmeridgian, above a poor ammonite record of rasenia sp. and at a level where bioturbated mudstones are followed by laminated mudstones, just as on store koldewey and in milne land. the appearance of p. capillosa and its associated assemblage may therefore reflect a relative sea-level rise. the presence of p. capillosa and its associated dinoflagellate cyst assemblage may be used as a general indication of the stratigraphical level from the r. cymodoce chronozone into the a. mutabilis (a. eudoxus?) chronozone. the few other associated species in the janusfjellet section are the stratigraphically long-ranging species paragonyaulacysta borealis, lanterna saturnalis and tubotuberella apatela, the typical borealis assemblage (brideaux & fisher 1976) which also occurs in east and north greenland. the dinoflagellate cyst assemblages in peary land and svalbard provide no clear upper stratigraphical limits, except for the disappearance of p. capillosa. p. capillosa occurs commonly up to the a. autissiodorensis chronozone (top kimmeridgian) in milne land but is also recorded scattered throughout the volgian. the successions at store koldewey, peary land and svalbard within the range of abundant p. capillosa are therefore considered to be of kimmeridgian (pre-volgian) age. discussion despite the similarity in composition and abundance of specific species in the geographically widespread assemblages described above, there are also clear differences reflecting the latitudinal distance between the compared localities. the upper jurassic dinoflagellate cyst assemblage on store koldewey is a mixture of species with subboreal or boreal preference but with a clear affinity to the boreal region. p. pannosum and o. patulum are abundant and long-ranging in northwest europe, they are abundant with a more restricted stratigraphical range in east greenland, and they appear in low numbers with a limited stratigraphical range in store koldewey. this is probably close to the northern limit of these species as they do not appear farther to the north in peary land (piasecki 1994) and svalbard (århus 1988). on store koldewey, the occurrence of these species so far north is associated with the most geus bulletin no 5.pmd 29-10-2004, 11:14109 110 significant relative sea-level rise recorded in the upper jurassic of east greenland, in the r. cymodoce to a. mutabilis (a. eudoxus) chronozones of the kimmeridgian. sequence stratigraphic implications studies of the jurassic ammonite and dinoflagellate cyst stratigraphy integrated with sedimentological studies and sequence stratigraphical interpretations in east greenland lead towards an integrated genetic model in which the units can be identified by their content of dinoflagellate cysts. the present study of the dinoflagellate cyst assemblages on store koldewey contributes to the study of this complex problem. the middle jurassic succession deposited directly on crystalline basement on store koldewey correlates with contemporaneous and stratigraphically similar successions from the lower part of the fossilbjerget formation in jameson land (p5 and p6 third-order sequences of engkilde & surlyk 2003), the charcot bugt formation in milne land (larsen et al. 2003) and the pelion formation at hold with hope (vosgerau et al. 2004, this volume). both in milne land and on hold with hope, the middle jurassic successions represent the earliest evidence of the jurassic transgression of the margins of the sedimentary basins. the ammonite fauna from the a. ishmae chronozone is among the most widespread faunas in the arctic (callomon 1993) and marks the considerable extent of this circum-arctic second-order marine transgression represented also in most areas of east greenland. the third-order depositional sequence p5 in jameson land is limited by sequence boundaries located in the a. ishmae and c. calyx chronozones, respectively (engkilde & surlyk 2003). the p6 depositional sequence is confined to the c. apertum, c. nordenskjoeldi and basal p. koenigi chronozones. stratigraphically, p5 and p6 correspond to the two sedimentological units identified in the middle jurassic of store koldewey and their content of dinoflagellate cysts correlates as well. in milne land, the charcot bugt formation also contains ammonites from the a. ishmae and the a. cranocephaloide zones (larsen et al. 2003). the associated dinoflagellate cyst assemblages correlate well with assemblages from the corresponding lower pelion succession on store koldewey. dinoflagellate cysts from the higher pelion formation, the spath plateau member, at store koldewey (c. apertum chronozone to the basal p. koenigi chronozone) correlate with those from assemblage 4 from the charcot bugt formation. on hold with hope, the pelion formation comprises two sedimentological units: a lower sandstone unit followed by the spath plateau member (vosgerau et al. 2004, this volume). ammonites indicating the c. apertum zone and the p. koenigi zone occur in the basal strata of spath plateau member. dinoflagellate cyst assemblages equivalent to the assemblages from the corresponding spath plateau member on store koldewey have been recorded in this succession (piasecki et al. 2004, this volume). thus, two distinct dinoflagellate cyst assemblages of bathonian and callovian age characterise the two sedimentological units that have been identified as third-order depositional sequences and are found to be extensively distributed throughout east greenland. as already mentioned above, the upper jurassic succession on store koldewey, consisting of the payer dal and bernbjerg formations, correlates excellently with contemporaneous successions from eastern peary land and svalbard in the north to milne land in the south. the overall transgressive trend from the upper oxfordian to maximum flooding in the kimmeridgian is represented by sedimentary deposits throughout east and north greenland that can be correlated in detail on the basis of both ammonites and dinoflagellate cysts. a succession of distinct dinoflagellate cyst assemblages characterises the stepwise progress of relative sea-level rise and can be recognised throughout the sedimentary basins of east greenland. conclusion the jurassic succession of store koldewey is divided into the pelion, payer dal and bernbjerg formations. abundant boreal ammonites date the succession in detail and associated floras of dinoflagellate cyst provide supplementary data. the pelion formation is dated as late bathonian – early callovian, in the middle jurassic, and the payer dal and bernbjerg formations are dated as late oxfordian – early kimmeridgian in the late jurassic. the middle jurassic succession consists of a bathonian and a callovian part. they can be correlated with contemporaneous sedimentary successions on hold with hope, in jameson land and in milne land to the south, where corresponding dinoflagellate cyst assemblages have been recorded. ranges of individual dinoflagellate species are given in relation to the boreal ammonite stratigraphy. geus bulletin no 5.pmd 29-10-2004, 11:14110 111 the upper jurassic succession is correlated with corresponding successions in hochstetter forland, hold with hope and milne land to the south and with those of peary land, north greenland, and svalbard to the north. passing northwards from east greenland via north greenland to svalbard, the upper jurassic dinoflagellate cyst floras show a transition from dominantly subboreal species in central east greenland to a distinct boreal flora in svalbard. the stratigraphic ranges of many species and their abundance decrease towards the north. in contrast, species with boreal affinity range southwards to milne land. the assemblages in store koldewey are transitional in composition. the sedimentological units of store koldewey are placed in the sequence stratigraphic framework developed for the jurassic in east greenland, and the associated dinoflagellate cyst assemblages are used to characterise and to identify these sequence stratigraphic elements. acknowledgements work began as part of the project ‘resources of the sedimentary basins of north and east greenland’, supported by the danish research councils. the work was completed with support from the carlsberg foundation (carlsbergfondet) ans. 980089/20-262. the authors are grateful to dr. a. wierzbowski and dr. g.f.w. herngreen for careful comments and constructive suggestions. references alsgaard, p.c., felt, v.l., vosgerau, h. & surlyk, f. 2003: the jurassic of kuhn ø, north-east greenland. in: ineson, j.r. & surlyk, f. 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(eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 893–930. milner, p.s. & piasecki, s. 1996: boreal middle jurassic dinoflagellate cyst stratigraphy of jameson land, east greenland. in: piasecki, s. et al. (eds): formation of source and reservoir rocks in a sequence stratigraphic framework, jameson land, east greenland. energy research programme efp-93, projects 1313/93-0010 and 0017. danmarks og grønlands geologiske undersøgelse rapport 1996/30, vol. i and ii, 46 pp. piasecki, s. 1980: middle to late jurassic dinoflagellate cyst stratigraphy from milne land (east greenland) correlated with ammonite stratigraphy, 167 pp. unpublished ph.d. thesis, university of copenhagen, denmark. piasecki, s. 1994: biostratigraphy of the jurassic – lower cretaceous ladegårdsåen formation, peary land. in: håkansson, e. (ed.): wandel sea basin: basin analysis. efp-91, project no. 0012, 1–14. piasecki, s. 1996: boreal dinoflagellate cyst stratigraphy of middle to upper jurassic sediments of milne land, east greenland. in: piasecki, s. et al. (eds): formation of source and reservoir rocks in a sequence stratigraphic framework, jameson land, east greenland. energy research programme efp-93, projects 1313/93-0010 and 0017. danmarks og grønlands geologiske undersøgelse rapport 1996/30, vol. i and ii, 100 pp. piasecki, s. 2001: three new middle jurassic dinoflagellate cysts of east greenland. neues jahrbuch für geologie und paläontologie. abhandlungen 219(1–2), 15–31. piasecki, s. & stemmerik, l. 2004: jurassic dinoflagellate cysts from hochstetter forland, north-east greenland. in: stemmerik, l. & stouge, s. (eds): the jurassic of north-east greenland. geological survey of denmark and greenland bulletin 5, 89–97 (this volume). piasecki, s., stemmerik, l., friderichsen, j.d. & higgins, a.k. 1994: stratigraphy of the post-caledonian sediments in the germania land area, north-east greenland. rapport grønlands geologiske undersøgelse 162, 183–190. piasecki, s., larsen, m., therkelsen, j. & vosgerau, h. 2004: jurassic dinoflagellate cyst stratigraphy of hold with hope, geus bulletin no 5.pmd 29-10-2004, 11:14111 112 north-east greenland. in: stemmerik, l. & stouge, s. (eds): the jurassic of north-east greenland. geological survey of denmark and greenland bulletin 5, 73–88 (this volume). ravn, j.p.j. 1911: on jurassic and cretaceous fossils from northeast greenland. meddelelser om grønland 45(10), 437–500. smelror, m. 1993: biogeography of bathonian to oxfordian (jurassic) dinoflagellates: arctic, nw europe and circum-mediterranean regions. palaeogeography, palaeoclimatology, palaeoecology 102, 121–160. stemmerik, l. & piasecki, s. 1990: post-caledonian sediments in north-east greenland between 76° and 78°30′n. rapport grønlands geologiske undersøgelse 148, 123–126. sykes, r.m. & callomon, j.h. 1979: the amoeboceras zonation of the boreal upper oxfordian. palaeontology 22(4), 839– 903. sykes, r.m. & surlyk, f. 1976: a revised ammonite zonation of the boreal oxfordian and its application in north-east greenland. lethaia 9, 421–436. vosgerau, h., larsen, m., piasecki, s., therkelsen, j. & surlyk, f. 2004: a new middle–upper jurassic succession on hold with hope, north-east greenland. in: stemmerik, l. & stouge, s. (eds): the jurassic of north-east greenland. geological survey of denmark and greenland bulletin 5, 51–71 (this volume). geus bulletin no 5.pmd 29-10-2004, 11:14112 geological survey of denmark and greenland bulletin 28, 2013, 73-76 73 titanium minerals in cameroon christian knudsen, joseph penaye, martin mehlsen, roger k. mclimans and feiko kalsbeek the mineral rutile (tio2) is a major ore of titanium, which is used in products such as white pigment and titanium metal. the global consumption of titanium minerals in 2011 was c. 6.7 million tonnes of which 0.7 million tonnes were rutile and 6 million tonnes ilmenite (tifeo3). rutile is almost pure tio2 and therefore more valuable than ilmenite (c. 1500 $/t and 300$/t, respectively). compared with ilmenite, rutile can be processed with lower consumption of chemicals and yields less waste products. rutile was mined in cameroon between 1935 and 1955 when a total of 15 000 tonnes of rutile were extracted from alluvial deposits. the french bureau de recherches géologiques et minières conducted a drilling programme in cameroon in the 1980s which identified c. 2.6 million tonnes of rutile in discontinuous occurrences with concentrations of c. 1%. most of the occurrences are located in smallto medium-sized riverbeds with a thickness of 1.5–4.5 m. the main alluvial rutile area is located around the town of akonolinga, 80 km east of yaoundé, the capital of cameroon (fig. 1). the rutile in the alluvial deposits was derived from the bedrock by weathering, and at some sites major, residual rutile deposits are reported from quaternary lateritic deposits. the geological survey of denmark and greenland conducted a project together with the institut de recherches géologiques et minières in cameroon to gain a better understanding of how rutile formed in the bedrock before it was weathered out and try to tie the rutile in the alluvial deposits to its source rocks. this is done by studying the compositional variation of the rutile in the alluvial deposits and comparing it with possible bedrock sources. the compositional variation of ilmenite and monazite ((la,ce)po4)) and the age distribution of zircon (zrsio4) in alluvial sand and bedrock were also investigated. the chemical compositions of minerals in the sediments are used to infer the bedrock source of the minerals. this has particular application in many areas of cameroon, such as the southern part of the country which is characterised by low relief and dense rain forest with bedrock outcrops that are sparse and difficult to find. cameroon is a country in west central africa (fig. 1) and is called ‘africa in miniature’ because of its cultural, geological and landscape diversity. the landscape includes beaches, deserts, mountains, rain forests and savanna. the highest point is the active volcano mount cameroon (4095 m), and the country is home for over 200 different linguistic groups with french and english as the official languages. compared with other african countries, cameroon is politically and socially stable. the country covers an area of 475 442 km2 with a population of c. 20 million of which 70% are christians and 20% muslims. geochronology of the yaoundé group stendal et al. (2006) suggested that the rutile in southern cameroon is associated with the neoproterozoic yaoundé group that primarily consists of garnet-bearing metamorphosed sediments such as shale and sandstone. these fig. 1. simplified geological map of cameroon. © 2013 geus. geological survey of denmark and greenland bulletin 28, 73–76. open access: www.geus.dk/publications/bull central african republic tchad nigeria gulf of guinea equatorial guinea gabon congo yaoundé tertiary to recent volcanics phanerozoic sediments neoproterozoic sediments neoproterozoic plutonic rocks neoproterozoic metamorphic volcanics yaoundé group neoproterozoic metasediments undifferentiated gneiss palaeoproterozoic gneiss archaean gneiss fig. 4 nyon g ri ver akonolinga 100 km 4°n national border 10°e africa http://en.wikipedia.org/wiki/central_africa http://en.wikipedia.org/wiki/mount_cameroon http://en.wikipedia.org/wiki/french_language http://en.wikipedia.org/wiki/english_language 7474 sediments are believed to have been deposited on the passive margin of the congo craton (nzenti et al. 1988), and available isotope data suggest that the sediments that formed the precursor of the yaoundé group were deposited during late neoproterozoic time. toteu et al. (2004) interpreted some of the sediments as deposited between 626 and 600 ma. in order to elucidate the timing of deposition and metamorphism of the yaoundé group, zircons from three samples of the yaoundé group have been dated by laser ablation inductively coupled plasma mass spectrometry at the geological survey of denmark and greenland (kalsbeek et al. 2013; for analytical procedures see frei et al. 2006). the results for one of the samples, a garnet-kyanite gneiss (ggu 512411), are shown in fig. 2. zircon grains from this sample have wide metamorphic rims which surround cores that are interpreted to represent remnants of originally detrital grains (inset in fig. 2). the rims yielded a weighted mean age of 635 ± 10 ma (n = 41, mswd = 0.48). most of the cores are of neoproterozoic age (mainly 650–1100 ma); a small number of palaeoproterozoic zircon grains are also present. in contrast to ggu 512411, two other samples from the yaoundé group (ggu 512401 and 512415) did not yield zircon ages younger than c. 950 ma. most zircon grains from these samples are of palaeoproterozoic age, but some mesoproterozoic and archaean grains are also present. the different detrital zircon age distributions indicate different source areas of the yaoundé group sediments. one source area may have been dominated by either palaeoproterozoic or neoproterozoic zircon grains, and another source area showed diverse age distribution patterns with archaean to neoproterozoic zircon grains. the rutile-forming event the metamorphic event that affected the sediments was related to the pan-african orogeny and the formation of the gondwana supercontinent. in cameroon, the pan-african orogeny was caused by the collision of the congo craton to the south with the nigerian shield to the north. in the areas where placer rutile deposits are found, the bedrock often consists of kyanite-bearing mica schist, indicating high-pressure, metamorphic conditions during this orogeny. rutile occurs together with garnet and kyanite and was formed by breakdown of titanium-bearing minerals such as ilmenite, biotite and muscovite. in cameroon, the rutile is commonly located within kyanite and garnet crystals (fig. 3) fig. 2. 207pb/206pb age distribution (sircombe 2004) of zircons from sample ggu 512411 from the yaoundé group. the inset shows a zircon grain with a distinct metamorphic rim. ages of cores and rims are shown separately; for simplicity no distinction is made between concordant and discordant analyses. 2 mm almandine quartz muscovite plag-na kyanite biotite rutile 33.78% 18.85% 17.35% 11.21% 10.02% 4.04% 2.07% 0 10 20 30 0 1 2 3 n um be r o f a na lys es core 1918 ± 29 ma rim 652 ± 42 ma 41 analyses of metamorphic rims mean age 635 ± 10 ma 72 analyses of detrital cores rims cores 207pb/206pb age (ga) fig. 3. mineral liberation analysis (mla) image of a kyanite garnet gneiss sample (ggu 512921) in which the rutile is enclosed in both garnet and kyanite; this indicates that the rutile was formed during prograde metamorphism – i.e. during increasing pressure and temperature. mla is a scanning electron microscope-based method where a polished, mounted part of the sample is divided into 10 × 10 µm pixels. the beam of electrons generates an x-ray spectrum, which is subsequently compared with a library of spectra representing different minerals. in this way, the instrument can recognise the minerals in the sample and generate an image of the mineralogy. the mla technique was described by fandrich et al. (2007). http://en.wikipedia.org/wiki/gondwana http://en.wikipedia.org/wiki/supercontinent 75 that formed during high-pressure metamorphism, and thus rutile must have formed during the same episode when pressure and temperature were rising (prograde metamorphism). however, it is possible that some of the rutile grains in the rocks of the yaoundé group are inherited from older rocks, as suggested by stendal et al. (2006). the temperature of rutile formation can be inferred from the rutile geothermometer (zack et al. 2004; tomkins et al. 2007). eighteen rutile-bearing rock samples from the yaoundé group were studied. as shown in fig. 4, the temperature during formation of the rutile varies significantly. the geographical distribution of the temperatures appears to show a pattern with a core area around yaoundé characterised by high temperatures in the 700–900°c range with areas to the north-east and south-west characterised by temperatures of 600–700°c. both east and west of the core area, the estimated temperatures of formation are lower, mainly in the range 500–700°c. it is not known if there was only one episode of rutile formation. geochronological work on the rutile is planned to test if it is possible to tie the formation of the rutile in the area to the pan-african orogeny. sediment sampling and analysis a programme combining collection of sediment samples from modern rivers and sampling of older river deposits was initiated in an area in the southern part of cameroon (figs 1, 4). the samples from streams were collected as heavy mineral concentrates, and the alluvial deposits were sampled using either a light auger driven by a small engine mounted on a tripod or a hand auger (fig. 5). the samples were analysed for major and trace elements using x-ray fluorescence and inductively coupled plasma mass spectrometry, respectively. the ratio between tio2 and fe2o3 was used to identify areas where the titanium is primarily located in rutile. the results show that areas with a tio2/fe2o3 ratio > 2 in stream sediments and alluvial deposits coincide with areas where yaoundé group sediments are found below the surficial deposits. in order to study the texture of the titanium mineral grains in stream sediments and alluvial sand, samples were subjected to ‘mineral liberation analysis’ (fig. 6). the analysis shows that the detrital rutile grains often have a rim of ilmenite. this is also seen in rock samples from the area fig. 4. pie diagrams showing temperatures during formation based on the zr content in rutile from 18 rock samples from the yaoundé group. the calculation of the temperature t (in °c) = 127.8 × ln (zr in rutile in ppm) –10 is based on zack et al. (2004). yg: yaoundé group. for location see fig. 1. fig. 5. two sampling methods being tested. left: a light auger driven by a small engine mounted on a tripod; right: a hand auger. at this location, the hand auger was superior as it penetrated 4.5 m in 30 minutes whereas the engine-driven auger penetrated 2 m during the same time. 12°e 3°n 100 km tertiary to recent volcanics phanerozoic sediments neoproterozoic sediments neoproterozoic plutonic rocks high-grade metamorphosed sediments of the yg low-grade metamorphosed sediments of the yg undifferentiated gneiss palaeoproterozoic gneiss archaean gneiss temperature (°c) legend to pies, 201–300 301–400 401–500 501–600 601–700 701–800 801–900 yaoundé 7676 where rutile is overgrown by ilmenite. this is caused by recrystallisation of the rock under lower-grade metamorphic conditions, probably retrogression during the late part of the pan-african orogeny. conclusions the titanium mineral rutile was formed by high-grade metamorphism of sedimentary rocks of the neoproterozoic yaoundé group during the pan-african orogeny c. 635 ma ago. the temperature of the rutile-forming event has been estimated using the ‘rutile geothermometer’ based on the zr content in the rutile, and it is found that the highest temperatures (c. 750 ± 100°c) are found in the area from yaoundé towards the north-east. both north-west and south-east of this area the rutile geothermometer indicates lower metamorphic temperatures (c. 600 ± 100°c). acknowledgements we thank j.v. hell, director of institut de recherches géologiques et minières in cameroon for enthusiastic support and permission to use vehicles from the institute during the field work. j. boserup constructed the small tripod drillrig and alfons berger assisted with the electron microprobe analyses at the department of geosciences and natural resource management, university of copenhagen. we thank j.z. johansen, j.m.u.o. njel and b. kankeu for help during field work in cameroon. dupont titanium technologies, wilmington, delaware, usa, provided financial support. references fandrich, r., gu, y., burrows, d. & moeller, k. 2007: modern sembased mineral liberation analysis. international journal of mineral processing 84, 310–320. frei, d., hollis, j.a., gerdes, a., harlov, d., karlsson, c., vasques, p., franz, c., johansson, l. & knudsen, c. 2006: advanced in situ geochronological and trace element microanalysis by laser ablation techniques. geological survey of denmark and greenland bulletin 10, 25–28. kalsbeek, f., ekwueme, b.n., penaye, j., de souza, z.s. & thrane, k. 2013: recognition of early and late neoproterozoic supracrustal units in west africa and north-east brazil from detrital zircon geochronology. precambrian research 226, 105–115. nzenti, j.p., barbey, p., macaudière, j. & soba, d. 1988: origin and evolution of the late precambrian high-grade yaoundé gneisses (cameroon). precambrian research 38, 91–109. sircombe, k.n. 2004: agedisplay: an excel workbook to evaluate and display univariate geochronological data using binned frequency histograms and probability density distributions. computers & geosciences 30, 21–31. stendal, h., toteu, s.f., frei, r., penaye, j., njel, u.o., bassahak, j., nni, j., kankeu, b., ngako, v. & hell, j.v. 2006: derivation of detrital rutile in the yaoundé region from the neoproterozoic pan-african belt in southern cameroon (central africa). journal of african earth sciences 44, 443–458. tomkins, h.s., powell, r. & ellis, d.j. 2007: the pressure dependence of the zirconium-in-rutile thermometer. journal of metamorphic geology 25, 703–713. toteu, s.f., penaye, j. & djomani, y.p. 2004: geodynamic evolution of the pan-african belt in central africa with special reference to cameroon. canadian journal of earth sciences 41, 73–85. zack, t., moraes, r. & kronz, a. 2004: temperature dependence of zr in rutile: empirical calibration of a rutile thermometer. contributions to mineralogy and petrology 148, 471–488. authors’ addresses c.k., f.k. & m.m., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ckn@geus.dk j.p., institut de recherches géologiques et minières, rue mgr vogt, p.o. box 4110 nlongkak, yaoundé, cameroon. r.k.mcl., dupont titanium technologies, experimental station, e352/217, route 141 and henry clay road, wilmington, de 19808, usa. 500 µm rutile ilmenite fe-oxides zircon leucoxene titanite silicates accessories pseudorutile titanomagnetite monazite fig. 6. mineral liberation analysis image of detrital grains from nyong river in cameroon (ggu 517615). the purple coloured parts are rutile and the blue is ilmenite. the orange grains are altered ilmenite grains (leucoxene). http://www.sciencedirect.com/science/article/pii/s0301926812003105 http://www.sciencedirect.com/science/journal/03019268 http://www.sciencedirect.com/science/journal/03019268/226/supp/c mailto:ckn@geus.dk geological survey of denmark and greenland bulletin 23, 2011, 69–72 69 2010, the year under review, marks the centennial of perhaps the most controversial structure in the arctic: the wegener fault, the 1000-km long fracture that is supposed to underlie nares strait and define the north-western margin of an independent greenland plate (fig. 1). the seaway between greenland and ellesmere island, canada, was branded a megashear by frank taylor who, purely on physiographic expression, postulated massive tertiary strike-slip (taylor 1910). this revolutionary idea fittingly found a place in alfred wegener’s theory of continental drift and thereafter in plate-tectonic theory with greenland drifting hundreds of kilometres from north america along what tuzo wilson subsequently dubbed the ‘wegener fault’ (wilson 1963). today, the concept lives on. in modern palaeogeography, nares strait is given a long multiphase dynamic history with collision of greenland and canada in the palaeogene (fig. 1). a freely drifting greenland plate unconstrained by ties to north america is now part of conventional wisdom as related in textbooks, review articles and educational material available on the internet. accordingly, the wegener fault is a standard feature in international compilations of world geology (e.g. unesco 2010; fig. 2). unfortunately, this 100-year acclamation from taylor (1910) to unesco (2010) is fundamentally flawed: the rocks and their relationships at nares strait flatly contradict the existence of the structure. scope and aim of this paper this paper’s four-page limit prevents discussion of the pros and cons of the wegener fault. for this, we refer to two multi-author volumes (dawes & kerr 1982; tessensohn et al. 2006) and to the latest papers (e.g. hansen et al. 2011; pulvertaft & dawes 2011). our aim is twofold: (1) to mark the centennial of a lithospheric structure disputed by the onsite geology, and (2) to add new evidence in the form of magnetic field variations across northern nares strait (kennedy channel) that define a lineament in harmony with previous interpretations of gravity data. regional setting the most recent overview of nares strait geology is by harrison et al. (2006). bedrock provinces of five ages are common to greenland and canada (fig. 2). in the south, archaean–paleoproterozoic crystalline shield is overlain by deposits of two sedimentary basins: the mesoproterozoic thule basin that straddles northern baffin bay and smith sound, and the e–w-trending palaeozoic franklinian basin that stretches westwards across northern canada into alaska. the cambrian–devonian fill of this basin is characterised by north-westerly thickening into a deep-water trough while its shelf overlaps the thule basin at smith sound. neoproterozoic basic dykes cut the shield kennedy channel and its geophysical lineaments: new evidence that the wegener fault is a myth thorkild m. rasmussen and peter r. dawes fig. 1. modern palaeogeography showing greenland as a wandering plate detached from canada throughout phanerozoic time: (a) cocks & torsvik (2006), (b–e) torsvik et al. in eide (2002), (f) present-day with the wegener fault (after taylor 1910; unesco 2010). red arrows mark the coast locations of magnetic and gravimetric anomalies described in this paper. 1000 km1000 km dc the 100-year myth: greenland as an independent drifting plate (1910–2010) 1000 km e 1000 km f 1000 km wegener fault present day 600 ma neoproterozoic 360 ma devonian 170 ma jurassic 80 ma cretaceous 60 ma paleocene ba 1000 km © geus, 2011. geological survey of denmark and greenland bulletin 23, 69–72. open access: www.geus.dk/publications/bull 7070 and its mesoproterozoic cover but are eroded off at the subcambrian (franklinian) unconformity. one e–w-trending basic dyke swarm has been mapped across southern kane basin and smith sound (fig. 2; oakey & damaske 2006). in end-devonian time, ellesmerian deformation transformed the franklinian trough into a fold belt flanked on the south by the homoclinal arctic platform. the fifth province – the carboniferous–cenozoic sverdrup basin – developed across the eroded, folded franklinian rocks but overlapped onto the platform. in palaeogene time, eurekan tectonism deformed the sverdrup basin and underlying franklinian rocks into a composite structural belt (innuitian orogen). geological relations at kennedy channel kennedy channel is sited within the franklinian basin with homoclinal strata of the arctic platform to the southeast and the folded trough to the north-west. nares strait trends roughly nne and thus oblique to the franklinian basin (fig. 2). however, the southern boundary of the folded trough has a sinuous form so that at kennedy channel structures roughly follow the coast, while to the south they swing westwards inland and to the north at hall basin, eastwards across greenland. limits of ellesmerian and eurekan diastrophism near the seaway roughly coincide producing a complicated fold-and-thrust belt of palaeozoic rocks with fault-bound packets of cretaceous–palaeogene deposits on ellesmere island (mayr 2008). homoclinal ordovician and silurian strata dipping 1–3° to the north-west underlie kennedy channel and these are involved in the eurekan fold-and-trust belt (harrison et al. 2007). in greenland, the cambro-silurian sedimentary pile overlying the shield is up to 3500 m thick and all evidence suggests that this geology continues offshore without structural break, with mid-channel hans ø and other islands exposing the uppermost reefal part of the silurian section (dawes 2004). we stress that the latest maps portray a homoclinal palaeozoic cover offshore unaffected by faulting (harrison et al. 2007). this contrasts with the thesis of some, for fig. 2. simplified geological map of the nares strait region. note that a submarine wegener fault, such as unesco (2010), must bypass obstacles like those across smith sound and kennedy channel (see fig. 4). fig. 3. regional, total magnetic intensity showing the same region as figs 2 and 4. reproduced from gaina et al. (2010). kennedy channel kane basin hall basin lincoln sea 200 km e l l e s m e r e i s l a n d d e v o n i s l a n d eureka n fo ld -a nd -t hr us t b el t inland ice greenland baffin bay arctic ocean nares strait wandel sea washington land hans ø smith sound carboniferous–cenozoic sverdrup and wandel sea basins folded trough arctic platform neoproterozoic basic dyke swarm mesoproterozoic thule basin archaean–palaeoproterozoic shield cambrian–devonian franklinian basin axel heiberg island 30°w 80°n 75°n 60°w 90°w 75°n 80°n g re en la n d c an ad a 1000 km wegener fault from unesco (2010) 399 287 231 194 164 141 121 105 90 77 64 52 41 31 21 11 2 –7 –15 –23 –31 –39 –46 –53 –61 –67 –74 –80 –87 –95 –104 –112 –120 –130 –142 –158 –182 –224 [nt] 250 km 120°w 86° 84° 82° 80° 78° 76° 80° 78° 76° 74° 90° 90°95°w 85° no data 80° 75° 70° 65°w 60° 30° 15°w 71 example, jackson et al. (2006, fig. 15), who draw a sinistral strike-slip fault just west of hans ø, a dislocation that “is hypothesized to be the leading edge of the plate boundary between the north american and greenland plates” (jackson et al. 2006, p. 21). its location in kennedy channel is roughly as shown in unesco (2010; fig. 2). the magnetic field anomalies our analysis is based on recently compiled magnetic and gravimetric data over the arctic (gaina et al. 2009; fig. 3). included are high-resolution data acquired by a canadian– german project that in 2001 used a helicopter from an icebreaker to survey kennedy channel (damaske & oakey 2006) and data acquired in 2003 by the geological survey of canada (oakey & damaske 2006). the data are low-pass filtered and levelled to provide a regional magnetic field representation of uniform spatial resolution. basically, a magnetic anomaly may be viewed as representing the response from a single isolated structure or the superimposed responses from several structures that merge into a well-defined anomaly. which approach is applicable is mainly controlled by distance between observation level and the structures, and by the cut-off wavelength in any applied lowpass filtering. our evaluation of continuity of anomalies is based on a tilt angle (miller & singh 1994) representation of a 5 km upward-continued version of the magnetic field. the upward continuation puts emphasis on structures having a regional extent but has the drawback of merging responses from adjacent structures. responses from shallow isolated structures are attenuated. the tilt angle provides structural information that is independent or unbiased with respect to magnetisation intensity of the structures. our focus here is on magnetic anomalies that can be linked to the crustal scale of kennedy channel close to the above-mentioned hypothetical plate boundary of jackson et al. (2006). relevant features are marked in fig. 4: a trend line based on peak values for the tilt angle and a previously published trend line of the nares strait gravity low (nsgl; oakey & stephenson 2008). we note that the magnetic trend line parallels the gravity trend for more than 1000 km along what is essentially the platform margin of the franklinian basin. this margin marks a drastic change in basin architecture from shallow shelf to deep trough with downwards flexuring of the substratrum or surface of the shield. the magnetic trend line is interrupted at kennedy channel but extrapolation along line provides an excellent match between the extended sections. the yellow dashed line crossing lincoln sea in fig. 4 represents merged responses from diverse unconnected magnetic structures seen both onshore and offshore on the total magnetic field map of damaske & oakey (2006, fig. 5) including responses from volcanogenic sandstones. depiction of these unconnected structures as a single anomaly is simply due to the above-mentioned lowpass filter properties of the applied upwards continuation. oakey & stephenson (2008) regard the nsgl to be an expression of low-density rocks within the franklinian basin and they demonstrate that the palaeogene eurekan frontal thrust (eft) obliquely truncates it. similarly, the magnetic trend is oblique to the eft. we interpret the magnetic anomalies paralleling the nsgl to reflect the lateral contrast in magnetic properties between franklinian basin strata and the crystalline shield. the continuity of both the magnetic and gravimetric trends across kennedy channel in the vicinity of hans ø implies that this area is not affected by a crustal dislocation. farther south at smith sound, e–w-trending, offshore, linear, magnetic anomalies represent dykes that correlate with neoproterozoic basic dykes onshore (oakey & damaske 2006). correlation between several offshore and onshore dykes of greenland is unequivocal, but on the opposfig. 4. tilt-angle map derived from 5 km upward-continued, total magnetic field from data in gaina et al. (2009). 1.47 1.37 1.28 1.18 1.08 0.98 0.88 0.78 0.69 0.59 0.49 0.39 0.29 0.20 0.10 0.00 –0.10 –0.20 –0.29 –0.39 –0.49 –0.59 –0.69 –0.79 –0.88 –0.98 –1.08 –1.18 –1.28 –1.37 –1.47 [radian] bouguer gravity trend line (nsgl) outlined by oakey & stephenson (2008) bouguer gravity trend line extended magnetic tilt trend line magnetic tilt trend line extrapolated lincoln sea magnetic tilt trend line no data 120°w 86° 84° 82° 80° 78° 76° 80° 78° 76° 74° 90° 90°95°w 85° 80° 75° 70° 65°w 60° 30° 15°w 250 km 7272 ing coast, although potential correlatives occur, there is a narrow coastal gap between magnetically identified offshore dykes and those onland (fig. 2; dawes 2009, fig. 5; pulvertaft & dawes 2011, fig. 3). we note here that on our magnetic tilt-angle map there is continuity in terms of texture of the magnetic tilt angle from the offshore area, intersected by dykes, to those onshore (fig. 4). conclusions a century after frank taylor’s proposal, some two dozen geological–geophysical markers within precambrian–palaeozoic rocks have been identified that demonstrate stratigraphic and structural continuity across nares strait (dawes & kerr 1982, pp. 369–386; tessensohn et al. 2006, pp. 129– 160). the magnetic lineament brought to notice here represents one more marker that militates against plate-boundary strike-slip deformation through the seaway. we take the persistence and parallelism of the magnetic and gravimetric anomalies to indicate crustal coherence between greenland and ellesmere island and we conclude that these geophysical lineaments are incompatible with the wegener fault (figs 1, 2). they confirm the story revealed by onshore outcrops that the franklinian basin is a structural entity stretching from ellesmere island to greenland unhindered by a lithospheric break. we challenge advocates of the 100-year model to explain how a major dislocation can be reconciled with the geophysical lineaments, as well as other obstacles that cross the waterway, for example, the neoproterozoic dyke swarm (fig. 2). furthermore, supporters of conventional reconstructions, such as those in fig. 1, must explain the plate-tectonic mechanisms by which such features are repositioned into perfect alignment (without offset), how basic dykes can preserve their linearity (without deformation) and how the harmonious within-plate pattern of the regional geology is reassembled (without mismatch). the kennedy channel geophysical lineaments – as well as two dozen previously defined markers – reflect intraplate geology that confirms the mythical character of the wegener fault which remains after a hundred years nothing more than a theory. references cocks, l.r.m. & torsvik, t.h. 2006: european geography in a global context from the vendian to the end of the palaeozoic. geological society of london memoir 32, 83–95. damaske, d. & oakey, g.n. 2006: volcanogenic sandstones as aeromagnetic markers on judge daly promontory and in robeson channel, northern nares strait. polarforschung 74, 9–19. dawes, p.r. 2004: explanatory notes to the geological map of greenland, 1:500 000, humboldt gletscher, sheet 6. geological survey of denmark and greenland map series 1, 48 pp. + map. dawes, p.r. 2009: precambrian–palaeozoic geology of smith sound, canada and greenland: key constraint to palaeogeographic reconstructions of northern laurentia and the north atlantic. terra nova 21, 1–13. dawes, p.r. & kerr, j.w. (eds) 1982: nares strait and the drift of greenland: a conflict in plate tectonics. meddelelser om grønland geoscience 8, 392 pp. eide, e.a. (coord.) 2002: batlas – mid norway plate reconstruction atlas with global and atlantic perspectives, 75 pp. trondheim: geological survey of norway. gaina, c., werner, s.c. & the camp-gm group 2009: circum-arctic mapping project – gravity and magnetic maps. geological survey of norway, report 2009.010, 21 pp. gaina, c., werner, s.c., mioara, m. & the camp-gm group 2010: magnetic and gravimetric anomaly maps of the arctic, 1: 5 000 000. unesco and geological survey of norway. hansen, k., dawes, p.r., frisch, t. & jensen, p.k. 2011: a fission track transect across nares strait (canada–greenland): further evidence that the wegener fault is a myth. canadian journal of earth sciences 48, 819–840. harrison, j.c., brent, t.a. & oakey, g.n. 2006: bedrock geology of the nares strait region of arctic canada and greenland, with explanatory text and gis content. geological survey of canada open file 5278, 60 pp. + map. harrison, j.c., dewing, k. & mayr, u. 2007: geology of hans island and adjacent parts of kennedy channel, northwest greenland (kalaallit nunaat) and northern nunavut (canada). geological survey of canada open file 5321, 33 pp. + map. jackson, h.r., hannon, t., neben, s., piepjohn, k. & brent, t.[a.] 2006: seismic reflection profiles from kane to hall basin, nares strait: evidence for faulting. polarforschung 74, 21–39. mayr, u. (ed.) 2008: geology of northeast ellesmere island adjacent to kane basin and kennedy channel, nunavut. geological survey of canada bulletin 592, 404 pp. miller, h.g. & singh, v. 1994: potential field tilt – a new concept for location of potential field sources. journal of applied geophysics 32, 213–217. oakey, g.n. & damaske, d. 2006: continuity of basement structures and dyke swarms in the kane basin region of central nares strait constrained by aeromagnetic data. polarforschung 74, 51–62. oakey, g.n. & stephenson, r. 2008: crustal structure of the innuitian region of arctic canada and greenland from gravity modelling: implications for the palaeogene eurekan orogen. geophysical journal international 173, 1039–1063. pulvertaft, t.c.r. & dawes, p.r. 2011: north atlantic spreading axes terminate in the continental cul-de-sacs of baffin bay and the laptev sea. canadian journal of earth sciences 48, 593–601. taylor, f.b. 1910: bearing on tertiary mountain belts and on the origin of the earth’s plan. bulletin of the geological society of america 21, 179– 226. tessensohn, f., jackson, h.r. & reid, i.d. (eds) 2006: nares strait and wegener transform fault. polarforschung 74, 198 pp. unesco 2010: geological map of the world, scale 1:25 000 000, 3rd edition. wilson, j.t. 1963: hypothesis of earth’s behaviour. nature 198, 925–929. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: tmr@geus.dk and prd@geus.dk geological survey of denmark and greenland. bulletin 10, 9-12 upper maastrichtian chalks form important hydrocarbon reservoirs in the danish sector of the north sea and have been intensively studied, yet their lithological uniformity can frustrate attempts to develop a high-resolution stratigraphic subdivision and a genetic understanding of the factors controlling production and sedimentation of the pelagic carbonate ooze. recent research into these topics, supported by the danish energy authority, was carried out by the geological survey of denmark and greenland (geus) in collaboration with the geological institute, university of copenhagen by means of a multidisciplinary study involving quantitative/semiquantitative palynology, micropalaeontology (nannofossils, foraminifers) and isotope geochemistry, integrated with detailed sedimentology. two key wells were selected, the m-10x well from the dan field and the e-5x well from the tyra se field (fig. 1), based on the extensive core coverage in these wells and on their position in the southern part of the danish central graben where evidence of large-scale resedimentation (and consequent stratigraphic complexity) is uncommon within the maastrichtian section. in focusing on such a pelagic carbonate system, the ultimate aim is a holistic understanding of the marine system including temperature variation, nutrient supply and distribution, salinity, watermass layering, circulation and oxygen distribution. all these factors influence organic productivity and thus the accumulation of biogenic sediment. this study concentrated on a number of palaeoceanographic signals that can be derived from the sedimentary record, summarised in fig. 2. planktonic organisms, both phytoplankton (e.g. coccolithophores, some dinoflagellates) and zooplankton (e.g. foraminifers) provide a record of conditions in the upper watermasses, largely within the photic zone, while bottom conditions are indicated by epifaunal/infaunal organisms (e.g. benthic foraminifers) and bioturbation, and by the sedimentological evidence of depositional processes at the sea floor. on a larger scale, the input of terrestrial organic material relative to the marine component can provide an indirect measure of shoreline migration and thus relative sea-level change, a factor that is also reflected in the δ13c isotopic composition of the seawater, as recorded by the biogenic carbonate ooze. © geus, 2006. geological survey of denmark and greenland bulletin 10, 9–12. available at: www.geus.dk/publications/bull stratigraphy and palaeoceanography of upper maastrichtian chalks, southern danish central graben jon r. ineson, bjørn buchardt, susanne lassen, jan a. rasmussen, poul schiøler, niels h. schovsbo, emma sheldon and finn surlyk 4°00' 5°00' 5°00' 55°30' 56°00' 56°30' east north sea block tyra-igor ridge salt dome province c offee soil fault tyra se field denmark dan field e-5x m-10x 25 km fault zone inversion zone oil field in chalk gas field in chalk studied wells uk norway denmark germany the netherlands 500 km fig. 1. map showing the location of the two key wells, m-10x (dan field) and e-5x (tyra se field), and the dominant late cretaceous structural elements in the danish central graben. 9 photic zone phytoplankton zooplankton benthic epi/infauna soft-bodied infauna seawater chemistry (13c, 18o) terrestrial organic matter sediment fabric/stru ctu re fig. 2. cartoon (scale arbitrary) showing the source of the palaeoceanographic signals utilised in the project; see text for explanation. stratigraphy a prerequisite for this palaeoceanographic study was a wellconstrained stratigraphic breakdown of the succession to permit confident intrabasinal correlation and comparison with regional/global events described in the literature. although awaiting formal assignment, the upper maastrichtian and danian chalks of the danish central graben are herein referred to the tor and ekofisk formations, respectively, as defined from the norwegian sector of the north sea. despite the numerous biostratigraphic studies undertaken on this stratigraphic interval in the danish sector, much of the information is in the form of company reports, and published data are few. a high-resolution biostratigraphic study was thus undertaken to identify internationally recognised biozones related to the dinoflagellate cyst, nannofossil and microfossil floras and faunas. detailed stable isotope investigations were incorporated to examine the potential role of chemostratigraphy. this resulted in a consistent framework for the uppermost maastrichtian of the two key wells in the southern danish central graben, exemplified by the subdivision of the m-10x cored section shown in fig. 3. as a supplement to the zonal biostratigraphy, semi-quantitative data provide a second tier of stratigraphic constraints; these include dinoflagellate (palaeocystodinium denticulatum, palynodinium grallator), foraminifer (praebulimina laevis, pseudotextularia elegans) and nannofossil (watznaueria barnesiae) acme events. the maastrichtian–danian boundary in the m-10x and e-5x cored sections is marked by a distinctive hardground developed in the uppermost c. 1 m of the tor formation. the association of complex cross-cutting thalassinoides networks, the cemented and bored upper layer and the irregular, pitted, glauconite-impregnated upper surface are indicative of a mature, evolved hardground surface, recording a protracted period of non-sedimentation at the sea floor. the occurrence of mixed late maastrichtian and early danian faunal and floral assemblages within the uppermost metre beneath the hardground surface suggests that the burrow systems remained open into danian time, thus accumulating complex multi-generational fills. the lowermost danian biozones are absent above the hardground indicating that this surface remained exposed at the sea floor for at least several hundred thousand years. palaeoceanography analysis of palaeoceanographic trends and evolution in the late maastrichtian in the study area has required integration of a wide range of detailed data, only a fraction of which can be reviewed in this summary. in the following, selected aspects of the full dataset (see ineson et al. 2004a) are presented to provide the essential framework of the palaeoceanographic model. sedimentology the upper maastrichtian chalk succession in the southern danish central graben is lithologically uniform, comprising pure coccolith-rich chalks with rare skeletal grains larger than silt grade, and a dominant ‘lime mudstone’ texture. in the dan field area, the cored upper maastrichtian section is divisible into a lower cyclic succession, characterised by metre-scale laminated/bioturbated cycles, and an upper succession that is thoroughly bioturbated and overtly non-cyclic (fig. 3). it has been proposed that the laminated facies in the 10 6618 6617 6616 6615 6614 6613 l it h o lo gy d ep th ( fe et ) d ep o si ti o n al / b io ge n ic st ru ct u re s fr ac tu re s, st yl o lit es cycle 2 cycle 3 cycle 1 6450 feet 6500 6550 6600 6650 b io tu rb at ed ch al k l am in at ed – b io tu rb at ed cy cl ic c h al k biostrat. l it h o st ra t. m ic ro . p al yn o . n an n o . c h ro n o st ra t. dan. ic o p d e h b o p gr fc s 2 3 a fc s 2 3 b u c 2 0 b u c 2 0 c u c 2 0 d p 1 a eko. fm t m a to r fo rm at io n t p e b ar re n n n t p 2 e * * u p p er m aa st ri ch ti an chalk mudstone enhanced lam/bio cyclicity chalk wackestone marly chalk hardground fig. 3. stratigraphic framework of the cored upper maastrichtian section in the m-10x well showing the broad subdivision into lower cyclic and upper non-cyclic, bioturbated chalks; the expanded section shows two laminated–bioturbated cycles, characteristic of the lower cyclic chalks. dan., danian; eko., ekofisk. palynological zones/subzones: ico, isabelidinium cooksoniae; pde, palaeocystodinium denticulatum; hbo, hystrichostrogylon borisii; pgr, palynodinium grallator; tma, tanyosphaeridium magdalium; tpe, thalassiphora pelagica. asterisk indicates mixed floral and faunal assemblages of late maastrichtian and danian aspect associated with the top tor formation hardground. lower chalks represents an alternation of thin (millimetrescale) chalk turbidites and hemipelagites; preservation of this primary sedimentary layering is thought to relate to low levels of oxygen on the sea floor (damholt & surlyk 2004). the alternation of laminated and bioturbated chalks thus records rhythmic shifts in bottom-water oxygenation; subtle shifts in the floral and faunal assemblages across these cycles suggest that such changes may have been controlled by variation in the degree of watermass stratification (ineson et al. 2004b). in addition to the lack of overt cyclicity and the ubiquitous bioturbation, the upper chalks are also lithologically more variable, including rare incipient hardground surfaces, skeletal-rich chalks (wackestones) and an isolated marly chalk bed (fig. 3). a subtle environmental change is thus indicated from a rhythmically stratified watermass to a more homogeneous watermass under which the sea floor was well ventilated and locally/periodically influenced by weak bottom currents. palynofacies and stable isotopes a number of palynological parameters can be utilised to provide a sea-level record, such as the relative proportions of marine and non-marine palynomorphs and the relative abundance of the genus impagidinium (an indicator of open ocean conditions) within the dinoflagellate assemblages. these parameters provide a consistent sea-level signal in the two central graben wells (ineson et al. 2004a). the lower, cyclic chalks record an overall gradual sea-level fall culminating in a ‘peak lowstand’ near the base of the p. grallator dinoflagellate zone (fig. 4, c. 6500 ft) coinciding with the level of the incipient hardground surfaces. palynofacies data from the succeeding bioturbated chalks record a subsequent sea-level rise and a variable signal up to the barren interval at the top of the maastrichtian section. it has been observed widely that the carbon isotope (δ13c) record broadly tracks the eustatic sea-level curve (e.g. jarvis et al. 2002). the δ13c values for the lower, cyclic chalks decrease gradually up-section, paralleling the palynofacies trend and adding support to the sea-level interpretation. in both wells, the lowest δ13c values in the upper maastrichtian section occur just beneath the base of the p. grallator zone. nannoflora and foraminifers the nannofloral assemblages are diverse (average species richness of 22) and show only minor compositional fluctuations with time, reflecting overall a stable, cool water (boreal), oligotrophic (low nutrient) setting. subtle shifts in the relative proportions of certain taxa, however, are thought to reflect changes in the nature of the upper watermasses, and although the precise control (i.e. temperature, nutrient level, salinity) can be difficult to isolate, interpretation can be facilitated by comparison with other datasets. two examples are shown in fig. 4. lucianorhabdus cayeuxii has been demonstrated to occur preferentially in relatively inshore shelf sediments (pospichal & wise 1990); the general increase in the relative abundance of this species in the upper half of the studied sec11 1.00 1.25 1.50 1.75 2.00 δ13c ‰ v-pdb log (marine/nonmarine ratio) -0.5 0.5 10-1 planktic/benthic ratio 0 % %0 50 0 50 20 40 60 80 100% 6450 feet 6500 6550 6600 6650 ic o p d e h b o p gr t m a t p e b ar re n d in o fla g. z o n e c h ro n o st ra t. d an . u p p er m aa st ri ch ti an lucianorhabdus cayeuxii high productivity indicators isotopic variation palynofacies nannoflora foraminiferal fauna fig. 4. selected data from the m-10x well, illustrating the palaeoceanographic evolution in the late maastrichtian, as discussed in the text. for stratigraphic terms and key, see fig. 3; the horizontal dashed line indicates the boundary between the lower cyclic and upper non-cyclic chalks. the nannofossil genera grouped together to provide an indication of changing fertility levels are biscutum, discorhabdus, chiastozygus and zeugrhabdotus. tion is thus compatible with the overall shallowing trend indicated by the sedimentological and palynological data. certain nannofossil species are considered to be indicative of increased nutrient levels (e.g. watkins 1989); the combined signal exhibited by these species clearly mirrors the subdivision of the succession into cyclic and non-cyclic chalks (fig. 4), possibly reflecting a change in nutrient partitioning in the water column under different palaeoceanographic states. the foraminifer assemblages are dominated by planktonic forms, particularly the ecologically tolerant heterohelix globulosa, and plankton/benthos (p/b) ratios are typically over 70% (fig. 4). the benthic fauna is of outer shelf aspect overall, with rare agglutinated forms. a feature of particular note in the m-10x foraminifer dataset is the anomalously low p/b values at 6490–6500 ft, associated with an increase in the relative abundance of a mid-shelf benthic foraminifer, p. laevis. this shallowing signal from the foraminifer data coincides with the sedimentological and palynofacies data indicating the ‘peak lowstand’ in the late maastrichtian chalk sea. acme of the planktonic foraminifer p. elegans and the nannofossil w. barnesiae in the upper, non-cyclic chalks are indicative of warmer surface waters in the region at this time and are symptomatic of a global warming event in the latest maastrichtian (huber & watkins 1992; olsen et al. 2001). late maastrichtian palaeoceanographic evolution palaeoecological data, from semi-quantitative analysis of foraminifer, coccolith and dinoflagellate faunas and floras, integrated with isotopic, palynofacies and sedimentological data has led to a model for the palaeoceanographic evolution of the danish central graben in the late maastrichtian involving two contrasting oceanographic systems. the lower half of the cored section records a cool-water, oligotrophic, deep shelf that was prone to stratification. breakdown of this rhythmically stratified system accompanied a long-term gradual fall in sea level, perhaps reflecting a depth-related threshold beyond which stratification was no longer favoured. the upper half of the cored section records an increasingly dynamic and varied mid-shelf to deep shelf setting with a complex blend of environmental signals – sea-level change, watermass warming and evidence of low but variable productivity. following turnover of the stratified system, the combined datasets record a progressive shallowing to a peak lowstand located just above the base of the p. grallator dinoflagellate zone boundary; this sea-level event can be correlated with a key sequence boundary in the type maastrichtian of the netherlands (schiøler et al. 1997). acknowledgement the danish energy authority (grant no.1313/01-0001) is thanked for financial support. references damholt, t. & surlyk, f. 2004: laminated–bioturbated cycles in maastrichtian chalk of the north sea: oxygenation fluctuations within the milankovitch frequency band. sedimentology 51, 1323–1342. huber, b.t. & watkins, d.k. 1992: biogeography of campanian–maastrichtian calcareous plankton in the region of the southern ocean: paleogeographic and paleoclimatic implications. in: kennett, j.p. & warnke, d.a. (eds): the antarctic paleoenvironment: a perspective on global change. antarctic research series 56, 31–60. ineson, j.r., buchardt, b., lassen, s., rasmussen, j.a., schovsbo, n.h., schiøler, p., sheldon, e. & surlyk, f. 2004a: palaeontology, stable isotopes and sedimentology of the upper maastrichtian, danish central graben: a record of palaeoclimatic and palaeoceanographic change. danmarks og grønlands geologiske undersøgelse rapport 2004/81, 20 pp. ineson, j.r., buchardt, b., lassen, s., rasmussen, j.a., schovsbo, n.h., schiøler, p., sheldon, e. & surlyk, f. 2004b: on the origin of laminated–bioturbated chalk cycles in the upper maastrichtian, danish central graben. danmarks og grønlands geologiske undersøgelse rapport 2004/82, 22 pp. jarvis, i., mabrouk, a., moody, r.t.j. & de cabrera, s. 2002: late cretaceous (campanian) carbon isotope events, sea-level change and correlation of the tethyan and boreal realms. palaeogeography, palaeoclimatology, palaeoecology 2948, 1–34. olsen, r.k., wright, j.d. & miller, k.g. 2001: paleobiogeography of pseudotextularia elegans during the latest maastrichtian global warming event. journal of foraminiferal research 31, 275–282. pospichal, j.j. & wise, s.w. 1990: maastrichtian calcareous nannofossil biostratigraphy of maud rise odp leg 113 sites 689 and 690, weddell sea. proceedings of the ocean drilling program, scientific results 113, 465–487. schiøler, p., brinkhuis, h., roncaglia, l. & wilson, g.j. 1997: dinoflagellate biostratigraphy and sequence stratigraphy in the type maastrichtian (late cretaceous), enci quarry, the netherlands. marine micropaleontology 31, 65–95. watkins, d.k. 1989: nannoplankton productivity fluctuations and rhythmically-bedded pelagic carbonates of the greenhorn limestone (upper cretaceous). palaeogeography, palaeoclimatology, palaeoecology 74, 75–86. authors’ addresses j.r.i., p.s., n.h.s. & e.s., geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ji@geus.dk b.b. & f.s., geological institute, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. s.l., 96 settrington road, london sw6 3ba, uk. j.a.r., geological museum, university of copenhagen, øster voldgade 5–7, dk-1350 copenhagen k, denmark. 12 geological survey of denmark and greenland bulletin 4, 2003, pp 85-88 85 jakobshavn isbræ (also known as sermeq kujalleq or ilulissat isbræ) is situated at about 69°10′n and 50°00′w in west greenland. this major outlet from the inland ice has an extremely high rate of movement (nearly 1 m/hour) and thus a high production of icebergs, which via the icefjord float westwards through disko bugt to davis strait (fig. 1). estimates of the iceberg production are in the range of 35 ± 10 km3 ice per year, more than 10% of the entire calf-ice production of the inland ice (e.g. bauer l968; bindschadler 1984). the icefjord into which sermeq kujalleq calves is kangia, best known in glaciological literature as jakobshavn isfjord. spectacular changes of the glacier were observed during 2002 and 2003 at the same time as it was nominated for inclusion in the unesco world heritage list under the name ‘ilulissat icefjord’. nomination of ‘ilulissat icefjord’ and sermeq kujalleq the eskimo ruins and archaeological sites in the region around the modern greenland township of ilulissat include representatives of all the cultural phases since the first eskimo settlement of greenland. the association with greenland’s most productive glacier makes the ‘ilulissat icefjord’ area a strong candidate for inclusion in the unesco world heritage list. in december 2000 the government of greenland decided to nominate ‘ilulissat icefjord’, and the geological survey of denmark and greenland (geus) was given the task of preparing the nomination document (fig. 1; mikkelsen & ingerslev 2003). a decision as to whether ‘ilulissat icefjord’ will be awarded world heritage list status jakobshavn isbræ,west greenland: the 2002–2003 collapse and nomination for the unesco world heritage list anker weidick, naja mikkelsen, christoph mayer and steffen podlech fig. 1. extent of the nominated area of ‘ilulissat icefjord’ (red boundary lines). the ablation area and front of the glacier sermeq kujalleq are included, together with the entire icefjord area. modified from mikkelsen & ingerslev (2003). geological survey of denmark and greenland bulletin 4, 85–88 (2004) © geus, 2004 will be made at the annual meeting of unesco in june 2004. the nomination document includes comprehensive documentation on observations and investigations of sermeq kujalleq. this major glacier, or ‘ice stream’, is situated in a subglacial trough that in its outer parts reaches depths of 1500 m below sea level (iken et al. 1993; clarke & echelmeyer 1996), and it can be traced inland for about 100 km (echelmeyer et al. 1991). general descriptions of the outer part of the glacier stress the quasistable conditions of the glacier front throughout the second half of the 20th century (stove et al. 1983; echelmeyer et al. 1991; sohn et al. l997a, b). however, during an inspection visit to the nominated area in august 2003 radical changes in the situation of the glacier front were found to have occurred compared to the description set out in the nomination document (fig. 2). prehistoric variations of sermeq kujalleq at the beginning of postglacial time (c. 9500 b.p.) the front of sermeq kujalleq was situated at the mouth of the fjord, about 50 km west of the front position shown in fig. 1, resting on a bank near ilulissat at depths of 200–300 m below present sea level (weidick 1994). subsequently, the ice margin retreated some 65–70 km, and at the end of the climatic optimum c. 5000 years ago the glacier front was located about 20 km east of the ice margin position of 1964 (weidick et al. 1990, fig. 4). the following neoglacial readvance culminated during the little ice age (a.d. 1500–1900), with the maximum position of the glacier front in 1851 (fig. 3; cf. bauer 1968). since the postglacial climatic optimum the ice margin is presumed to have advanced in pulses, such as it has been documented for the ice margin 40 km north of sermeq kujalleq where the response of the ice-sheet margin has been calculated for the last 1400 years (reeh 1983). there may have been two periods of advance: at a.d. 700–800 and during the little ice age (a.d. 1500–1900). historical records of sermeq kujalleq historical records of the glacier fluctuations of sermeq kujalleq have been collected and described by larsen & meldgaard (1958) and georgi (1960). the numerous observations since the beginning of the 1700s and up to 1851 suggest a gradual advance during this period. subsequent to 1851 observations and descriptions are more frequent and more detailed. the first mapping of the frontal position was by rink in 1851 (rink 1857), the first determination of the fast movement of the glacier front was made in 1875 (helland 1876), and seasonal fluctuations of the glacier front position were recorded by investigations in 1879–80 (hammer 1883). regular meteorological observations were begun in the town of ilulissat in 1873. the quasistable period of sermeq kujalleq 1950–2002 recessional positions of sermeq kujalleq are summarised in fig. 3. the recent part of the curve is based on aerial photographs (from about 1950) and since 1962 also satellite information (sohn et al. 1997a). this well-documented 5086 fig. 2. sermeq kujalleq on 28 may 2003; glacierfront indicated by dashed red line. satel-lite image by aster (advanced spaceborne thermal emission and reflection radio-meter) installed in the terra satellite, with the position of the glacier front on 7 july 2001 indicated by dashed black line. aster data are distributed by the land processes distributed active archive center (lpdaac), located at the united states geological survey’s eros data center: http://lpdaac.usgs.gov year period coincides with a stable position of the glacier front at a broad part of the fjord. the quasistability of the glacier front position was probably influenced by subglacial topography (echelmeyer et al. l991; weidick 1992). the thinning of the glacier during the recessional period c. 1850–1950 has been estimated at more than 200 m (weidick 1992), while observations on subsequent changes in thickness of the glacier front are few and scattered. a lowering of the frontal surface may have occurred from the 1960s to the 1980s (echelmeyer et al. 1991). between 1993/94 and 1998/99 investigations of the glacier by laser altimetry showed sermeq kujalleq to be one of the few greenland outlets showing signs of slight growth (abdalati et al. 2001). however, since 1997 a sudden transition to a rapid thinning has occurred, starting in the lower reaches of the glacier and spreading gradually inland. by 2001 almost the entire glacier up to elevations of 2000 m exhibited thinning (thomas et al. 2003). substantial changes of the glacier front were therefore not entirely unexpected. the break-up of the sermeq kujalleq front the inhabitants of ilulissat had observed an unusual amount of detached parts of the glacier front at the mouth of the icefjord in 2002, and observations during the 2003 visit showed that these parts are the result of a break-up of the floating glacier front. to date more closely the time of glacier break-up and retreat of the front, studies were made of a series of landsat images covering the years 2001 to early 2003. it was established that the last winter advance of the quasistable period took place in march 2002, and was followed by a period of continuous break-up and recession of the glacier front. by may 2003, the glacier front was situated c. 11 km east of the ‘normal’ winter position. major parts of the floating glacier described by echelmeyer et al. (199l) had disappeared. two stages in this break-up are illustrated in fig. 2 (july 2001 and may 2003). the ‘normal’ retracted summer position of the quasistable stage (july 2001) was followed by the last winter advance of c. 3.5 km to march 2002. in may 2003, a major recession of the glacier front has occurred, and the glacier segment in tissarissoq had become isolated and had partially disintegrated. the frontal position after 7 july 2003 is currently uncertain, but recession seems to have continued and it is now approaching the retracted position of the climatic optimum 4000–5000 years ago. the occurrences of marine shells in the neoglacial moraines surrounding tissarissoq demonstrate that the area was ice-free in the past, although the duration of this ice-free period is not known. a single radiocarbon dating 87 fig. 3. top: approximate recessional positions of sermeq kujalleq. modified from bauer (1968). bottom: conceptual recessional curve of the sermeq kujalleq glacier front, based on the positions given by bauer (1968) up to 1964. the younger parts of the curve are based on satellite information by stove et al. (1983), sohn et al. (1997a, b) and later landsat and aster images. the width of the curve depicts range of seasonal variations in the position of the glacier front. note the rapid break-up and retreat from 2002–2003. bauer (1968) and georgi (1960) give very advanced positions of the front in 1870 and 1880; both were considered uncertain by engell (1904), and have therefore been omitted here. 88 of a walrus tusk is currently available (4290 ± 100 b.p.; weidick 1992). local legends record that this area was formerly ice-free and used as a hunting locality (hammer 1883, p. 5). if these legends record a real event, then the end of the open-water period and advance of the ice may be as recent as the little ice age (a.d. 1500–1900). this can only be verified by extensive dating of the marine deposits, such as that carried out to determine the duration of open water in nioghalvfjerdsfjorden in north-east greenland (bennike & weidick 2001). acknowledgement the nomination project received financial support from the danish ministry of the environment as part of the environmental support programme dancea – danish co-operation for environment in the arctic. references abdalati, w., krabill, w., frederick, e., manizade, s., martin, c., sonntag, j., swift, r., thomas, r., wright, w. & yungel, j. 2001: outlet glacier and marginal elevation changes: near-coastal thinning of the greenland ice sheet. journal of geophysical research 106(d24), 33729–33741. bauer, a. 1968: missions aériennes de reconnaissance au groenland 1957–1958. observations aériennes et terrestres, exploitation des photographies aériennes, détermination des vitesses des glaciers vêlants dans disko bugt et umanak fjord. by a. bauer, in collaboration with m. baussart, m. carbonnell, p. kasser, p. perroud and a. renaud. meddelelser om grønland 173(3), 116 pp. bennike, o. & weidick, a. 2001: late quaternary history around nioghalvfjerdsfjorden and jøkelbugten, north-east greenland. boreas 30, 205–227. bindschadler, r.a. 1984: jakobshavn glacier drainage basin: a balance assessment. journal of geophysical research 89, 2066–2072. clarke, t.s. & echelmeyer, k. 1996: seismic-reflection evidence for a deep subglacial trough beneath jakobshavns isbræ, west greenland. journal of glaciology 42(141), 219–232. echelmeyer, k., clarke, t.s. & harrison, w.d. 1991: surficial glaciology of jakobshavns isbræ, west greenland. part i: surface morphology. journal of glaciology 37(127), 368–382. engell, m.c. 1904: undersøgelser og opmaalinger ved jakobshavns isfjord og orpigsuit i sommeren 1902. meddelelser om grønland 26(1), 70 pp. georgi, j. 1960: otto fabricius und andere über die eisverhältnisse auf grönland, mit einem exkurs auf den jakobshavner gletscher. polarforschung 4 (jahrgang 28, 1958, heft 1–2), 79–91. hammer, r.r.j. 1883: undersøgelser ved jakobshavns isfjord og nærme ste omegn i vinteren 1879–1880. meddelelser om grønland 4(1), 68 pp. helland, a. 1876: om de isfyldte fjorde og de glaciale dannelser i nordgrønland. archiv for matematik og naturvidenskab 1, 69 pp. iken, a., echelmeyer, k. & funk, m. 1993: mechanism of fast flow in jakobshavn isbræ, west greenland. part i: measurements of temperature and water level in deep bore holes. journal of glaciology 39, 15–25. larsen, h. & meldgaard, j. 1958: paleo-eskimo cultures in disko bugt, west greenland. meddelelser om grønland 161(2), 75 pp. mikkelsen, n. & ingerslev, t. (eds) 2003: nomination of the ilulissat icefjord for inclusion in the world heritage list. document prepared for unesco, 136 pp. copenhagen: geological survey of denmark and greenland. reeh, n. 1983: ikke-stationær beregningsmodel for indlandsisens randzone. gletscher-hydrologiske meddelelser grønlands geologiske undersøgelse 83/7, 81 pp. rink, h. 1857: grønland, geografisk og statistisk beskrevet. vol. 1: det nordre inspektorat, 420 pp.; vol. 2: det søndre inspektorat, 588 pp. københavn: andr. fred. høst. sohn, h.g., jezek, k.c. & van der veen, c.j. 1997a: jakobshavn glacier, west greenland: 30 years of space-borne observations. geophysical research letters 25, 2699–2702. sohn, h.g., jezek, k.c. & van der veen, c.j. 1997b: seasonal variations in terminus position of jakobshavn glacier, west greenland. in: van der veen, c.j. (ed.): calving glaciers. byrd polar research center report 15, 137–140. stove, g.c., green, k., birnie, r.v., davidson, g., bagot, k., palmer, m., kearn, g., ritchie, p.f.s. & sugden, d.e. 1983: monitoring iceberg production from west greenland tidewater glaciers using landsat data. results of the agrispine experiment for the jakobshavn isbræ, 32 pp. macaulay institute for soil research, craigiebuckler, aberdeen, scotland and the national remote sensing centre, farnborough, hampshire, england. thomas, r.h., abdalati, w., frederick, e., krabill, w., manizade, s. & steffen, k. 2003: investigation of surface melting and dynamic thinning on jakobshavn isbræ, greenland. journal of glaciology 49(165), 231–239. weidick, a. 1992: jakobshavn isbræ area during the climatic optimum. rapport grønlands geologiske undersøgelse 155, 66–72. weidick, a. 1994: fluctuations of west greenland calving glaciers. in: reeh, n. (ed.): workshop on the calving rate of west greenland glaciers in response to climate change, 143–168. copenhagen, denmark: danish polar center. weidick, a., oerter, h., reeh, n., thomsen, h.h. & thorning, l. 1990: the recession of the inland ice margin during the holocene climatic optimum in the jakobshavn isfjord area of west greenland. palaeogeography, palaeoclimatology, palaeoecology 82, 389–399. authors’ addresses a.w., n.m. & c.m., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: awe@geus.dk s.p., department of geophysics, university of copenhagen, juliane maries vej 30, dk-2100 copenhagen ø, denmark. << /ascii85encodepages false /allowtransparency false /autopositionepsfiles true /autorotatepages /all /binding /left /calgrayprofile (dot gain 20%) /calrgbprofile (srgb iec61966-2.1) /calcmykprofile (u.s. web coated \050swop\051 v2) /srgbprofile (srgb iec61966-2.1) /cannotembedfontpolicy /warning /compatibilitylevel 1.4 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<feff4f7f752890194e9b8a2d5b9a5efa7acb76840020005000440046002065874ef65305542b8f039ad876845f7150cf89e367905ea6ff0c4fbf65bc63d066075217537054c18cea3002005000440046002065874ef653ef4ee54f7f75280020004100630072006f0062006100740020548c002000520065006100640065007200200035002e0030002053ca66f465b07248672c4f86958b555f3002> >> >> setdistillerparams << /hwresolution [2400 2400] /pagesize [595.000 842.000] >> setpagedevice geological survey of denmark and greenland bulletin 35, 2016, 63-66 63© 2016 geus. geological survey of denmark and greenland bulletin 35, 63–66. open access: www.geus.dk/publications/bull in the rinkian belt of west greenland, reworked archaean gneisses are overlain by supracrustal successions of the palaeoproterozoic karrat group, defined by henderson & pulvertaft (1967) as comprising two formations: the qeqertarssuaq formation and the nûkavsak formation. the group was later extended to include the mârmorilik formation (henderson & pulvertaft 1987) originally considered to be of archaean age (henderson & pulvertaft 1967) but later shown to be palaeoproterozoic, resting with a depositional unconformity on archaean gneiss (garde 1978). henderson & pulvertaft (1987) suggested that the carbonatedominated mârmorilik formation in the south was laterally equivalent to the siliciclastic-dominated qeqertarssuaq formation in the north, the two subbasins being separated by a basement topographic high. the karrat group and the archaean basement were metamorphosed and folded during the rinkian orogeny (henderson & pulvertaft 1987; grocott & pulvertaft 1990). during shipand helicopter-supported fieldwork in 2015, the kangerluarsuk–maarmorilik area (fig. 1) was visited to sample the zn-pb-mineralised horizons found by rtz mining and exploration ltd. (coppard et al. 1992) along the basement-cover boundary and to obtain oblique photographs of the contact that was described as faults on the mârmorilik geological map sheet (scale 1:100  000, henderson & pulvertaft 1987). the aim of this paper is to show preliminary new results on the geometry and type of unconformity between the palaeoproterozoic karrat group and the underlying gneisses, and to describe a new stratigraphic unit representing here the lowermost sedimentary sequence above the unconformity. this unit is of particular relevance as it controls the distribution of the zn-pb mineralisation. palaeoproterozoic palaeovalleys as a result of unconformity one of the best localities to characterise the basal unconformity of the palaeoproterozoic karrat group is along a cliff face of qaarsukassak at the head of kangerluarsuk (fig. 1). the more than 1800 m high cliff face exposes archaean basement rocks (umanak gneiss) and greywackes of the nûkavsak formation (henderson & pulvertaft 1987). the contact between the archaean basement and the karrat group was originally mapped as a series of faults, probably based on binocular observations from a boat in the 1960s, and from the black and white aerial photographs that were available at that time. the detail of our new oblique photographs clearly shows that this contact corresponds to a nonconformity with incised valleys into archaean basement gneiss filled by palaeoproterozoic siliciclastic rocks of the karrat group (figs 2, 3). the palaeotopography that is preserved below the palaeoproterozoic sedimentary cover appears deeply excavated with highs and lows representing an ancient palaeodrainage system with a topographic relief of 300–400 m (fig. 3). qaarsukassak represents one of the domal structures described by henderson & pulvertaft (1987), and the palaeotopography is accentuated by this fold structure. the south-eastern and north-western limbs of this wide upright anticline were visited and the basal unit was sampled (fig. 2). palaeovalleys at the basal unconformity of the palaeoproterozoic karrat group, west greenland pierpaolo guarnieri, camille a. partin and diogo rosa tornit 'discovery’ qaarsukassak alfred wegener halvø maarmorilik 51°30' w 52° w 7 1 °2 0 ' n 7 1 °1 0 ' n tornit ice nûkavsak fm metavolcanic rocks quaternary qaarsukassak fm umanak gneiss mârmorilik fm figs 2, 3 'discovery’ qaarsukassak alfred wegener halvø maarmorilik 51°30' w 52° w 7 1 °2 0 ' n 7 1 °1 0 ' n 5 km kangerluarsuup sermia ka nge rlu ars uk k greenland fig. 1. simplified geological map of the central karrat group area (modified from escher & pulvertaft 1995). k: kangerlussuakassak. 6464 the overall geometry of the archaean–palaeoproterozoic unconformity displays important regional differences. in the maarmorilik sector to the south the palaeosurface appears to be a peneplain covered by the basal quartzites of the mârmorilik formation described by garde (1978); in the central kangerluarsuk sector described here, it is represented by well-developed incisions filled with a new stratigraphic unit and local occurrences of lava flows. to the north of the studied area, in the karrat isfjord sector, the unconformity is represented by an undulated palaeosurface covered by thick metavolcanic rocks that were also noted by henderson & pulvertaft (1987). qaarsukassak formation here we describe a new stratigraphic unit in the karrat group, informally named the qaarsukassak formation, which occurs locally below the nûkavsak formation in the kangerluarsuk fjord area (fig. 1). this sequence was previously described at the so-called ‘discovery’ showing by coppard et al. (1992). the report describes a 30–66 m thick quartzite-carbonate succession with a mineralised zone occurring dominantly in calcite-bearing dolostone, and bounded by the archaean basement gneiss and the nûkavsak formation. the measured section from the base to the inferred stratigraphic top (fig. 4) reveals that its stratigraphic thickness prior to structural repetition is less than 20 m. the lower contact with the archaean basement gneiss is a planar to undulating erosional surface that preserves a depositional contact. laminated to massive quartzite fines upward into fine-grained metamorphosed sandstones and sandy mudstones, including calcite-cemented and graphitic quartzites (unit 1). these siliciclastic rocks are overlain in sharp contact by light grey to white metacarbonate rocks with pods of massive tremolite and in some horizons, minor graphite (unit 2). this is succeeded by another quartzite unit (unit 3), followed by dark grey, laminated metacarbonate rocks with possible slump folds and minor tremolite veining (unit 4). the overlying, rusty weathering metasedimentary rocks including graphitic, metamorphosed mudstones and siliciclastic rocks represent the ore zone (unit 5). the upper contact of the qaarsukassak formation with the nûkavsak formation is not well-exposed. also of note is a thin re-sedimented calcitic marble horizon within the basal part of the nûkavsak formation, which might be derived from contemporaneous erosion of the qaarsukassak formation. all metacarbonate rocks in the qaarsukassak formation at the ‘discovery’ section described here are calcitic. the mârmorilik formation, by contrast, contains both calcitic and dolomitic marble (garde 1978). the primary occurrence of calcitic marble in the upper mârmorilik formation suggests possible correlation with the qaarsukassak formation. the presence of tremolite in some outcrops of qaarsukassak kangerlussuakassak 1850 m nw nukavsak fm (palaeoproterozoic) qaarsukassak fm (palaeoproterozoic) umanak gneiss (archaean) oblique photo (fig. 3) se 0 m 1000 2000 ‘discovery’ 569813 569816 fig. 2. geological cross-section of qaarsukasak showing the palaeoproterozoic unconformity with the incised palaeovalleys into the archaean umanak gneiss infilled by the qaarsukassaq formation at the bottom and the nûkavsak formation on top. the locations along the cross-section of two analysed samples (fig. 5) are also shown. qaarsukassak 1710 m 1140 m 960 m 1200 m nûkavsak fm umanak gneiss nw se u nconformity u nconformity fig. 3. the geometry of valleys and onlap of palaeoproterozoic greywackes of the nûkavsak formation at qaarsukassak, photographed from a helicopter. the difference between bottom and top of the palaeorelief is about 240 m. position of image shown on fig. 2. 65 the qaarsukassak formation, though, suggests the former presence of dolomite or high-mg calcite. unit 1 of the qaarsukassak formation could represent a fluvial environment, but outcrops lack the sedimentary structures to confirm this. instead, the qaarsukassak formation was likely deposited in a shallow marine environment. the qaarsukassak formation shows minor thickness variations along strike, suggesting that its deposition infilled pre-existing topographical lows in the archaean basement gneiss. in addition to the main section in the ‘discovery’ area, similar rocks also occur at two other localities, namely at tornit on the south side of kangerluarsuk and along kangerluarsuup sermia (fig. 1), although exposures are not laterally continuous. at tornit the orientation of the qaarsukassak formation is vertical and forms part of the overturned limb of the kigarsima nappe (henderson & pulvertaft 1987). at this outcrop the formation occurs between a thin amphibolite unit and the nûkavsak formation as a c. 10 m thick section comprised of calcitic marble and rusty metasedimentary rocks, with quartzite, a possible quartz-pebble conglomerate and siliceous marble. the section might be repeated by folding and therefore only less than 10 m thick. at kangerluarsuup sermia an overturned 30–40 m section of dark grey, tremolite-bearing dolomitic marble with black layers of metamudstone occurs structurally below archaean basement gneiss, which is thrust over the marble. the qaarsukassak formation extends to the north-west side of the section shown in fig. 2 on the northwestern limb of the wide anticline at kangerlussuakassak. mineralisation carbonate-hosted zn-pb mineralisation is common in the calcitic marble of the upper mârmorilik formation and was exploited at the black angel mine between 1973 and 1990 (thomassen 1991). additionally, coppard et al. (1992) identified intermittent stratabound zn-pb mineralisation of the ‘discovery’ and kangerluarsuup sermia occurrences, hosted within a carbonate and siliciclastic sequence. this sequence was considered to be the basal sequence of the mârmorilik formation by coppard et al. (1992). here, we allocate this succession to the qaarsukassak formation, since it is not in contact with the mârmorilik formation and the two units are separated by the basement topographic high. the recognition and understanding of the distribution of this formation are important as the formation controls the distribution of the regional zn-pb mineralisation, which should be distinguished from rusty* ** * * * * * * * * * 0 5 m 10 m 15 m 20 m unit 1 unit 2 unit 3 unit 4 ore zone u m an ak gn ei ss q aa rs u ka ss ak f m massive to laminated quartzite unconformity orthogneiss calcite-cemented quartzite graphitic quartzite metapelite carbonate rock pb-zn ore zone tremolite in carbonate rock slump folds palaeoproterozoic qaarsukassak fm archaean * fig. 4. stratigraphic section of the qaarsukassak formation at the type locality of qaarsukassak at head of kangerluarsuk (fig. 2 for location). fe 10(zn+pb) s sph py gn base metal mineralisation, hosted by carbonates (mârmorilik and qaarsukassak fms) barren, pyritic and/or pyrrhotitic mineralisation (nûkavsak fm) m et as ed im en t, py ri tic c he rt /s ha le g os sa ns 569807 569813 569816 fig. 5. fe–s–10(zn+pb) ternary diagram with discriminant compositional areas of base-metal mineralisation hosted by carbonates of the mârmorilik and qaarsukassak formations (right) and the rusty-weathering zones hosted by the nûkavsak formation (left). py: pyrite. gn: galena. sph: sphalerite. 6666 weathering, pyritic or pyrrhotitic, but barren horizons located in the nûkavsak formation. a ternary diagram (fig. 5) discriminates the signatures of the base-metal mineralisation hosted by carbonate rocks of the mârmorilik or qaarsukassak formations, and of the rusty-weathering horizons hosted by the nûkavsak formation. the latter are dominated by pyrite or pyrrhotite and lack significant amounts of base metals (fig. 5). due to the effects of weathering, resulting in gossanous horizons, compositions of the sulphides in exposed rocks can depart from pyrite and become enriched in iron and depleted in sulphur. the former have variable amounts of iron sulphide, but can also have significant sphalerite or galena contents. the two samples closest to the metal corner were collected at the previously established ‘discovery’ and ‘south lakes glacier’ occurrences, samples 569813 and 569807, respectively. a third sample (569816, fig. 5) was collected c. 10 km to the north-west (fig. 2) where mineralisation was not previously seen. although the absolute base-metal content of this sample is not high, it is elevated relative to sulphur (akin to the tenor concept in nickel exploration). its presence supports potential correlation with the mârmorilik formation and the proposed continuation of the qaarsukassak formation, hosting base-metal mineralisation, to the north-western limb of the wide anticline of the cross section shown in fig. 2. discussion and conclusion henderson & pulvertaft (1987) proposed a palaeoproterozoic structural high of archaean basement gneiss exposed on alfred wegener halvø, bounding the carbonate-dominated mârmorilik basin to the south and separating it from the qeqertarssuaq and nûkavsak formations in the north. our new observations show that such archaean basement gneiss in the kangerluarsuk area is heavily incised, and that palaeoproterozoic palaeovalleys were filled in with siliciclastic and carbonate rocks of the qaarsukassak formation (new informal unit) that correlates well with the mârmorilik formation and is conformably overlain by the nûkavsak formation. its intermediary, mixed siliciclasticcarbonate sedimentation could have been marginal marine and transitional into sedimentation of the turbiditic greywackes of the nûkavsak formation. the palaeovalleys observed at qaarsukassak seem to reflect an erosional event prior to the deposition of a sedimentary cycle represented by a carbonate-rich basin in the maarmorilik area and an area rich in volcanic components in the north, separated by a continental shelf area. the erosion may have been related to a re-organisation of the basin structure, where uplift in some areas created the observed palaeotopography beneath the nûkavsak formation. the observations in the kangerluarsuk area described here, coupled with new observations from the qeqertarssuaq formation in the north which will be described elsewhere, may suggest that the karrat group comprises more than one sedimentary cycle separated by erosion and possibly deformation. an ongoing geochronological study of detrital zircon from different units within the karrat group is expected to yield new information about the depositional ages of its individual formations. acknowledgments this work was carried out within the framework of an ongoing project financed by the geological survey of denmark and greenland (geus) and the ministry of mineral resources of greenland (mmr). references coppard, j., swatton, s. & harris, c.j. 1992: karrat exclusive exploration licence. 1992 year-end report, 19 pp. unpublished report, rtz mining and exploration limited (in archives of the geological survey of denmark and greenland, geus report file 21297). escher, j.c. & pulvertaft, t.c.r. 1995: geological map of greenland, 1:2 500 000. copenhagen: geological survey of greenland. garde, a.a. 1978: the lower proterozoic marmorilik formation, east of mârmorilik, west greenland. meddelelser om grønland 200(3), 71 pp. grocott, j. & pulvertaft, t.c.r. 1990: the early proterozoic rinkian belt of central west greenland. in: lewry, j.f. & stauffer, m.r. (eds) the early proterozoic trans-hudson orogen of north america. geological association of canada special paper 37, 443–463. henderson, g. & pulvertaft, t.c.r. 1967: the stratigraphy and structure of the precambrian rocks of the umanak area, west greenland. meddelelser fra dansk geologisk forening 17, 1–22. henderson, g. & pulvertaft, t.c.r. 1987: geological map of greenland, 1:100 000. mârmorilik 71 v.2 syd, nûgâtsiaq 71 v.2 nord, pangnertôq 72 v.2 syd. lithostratigraphy and structure of a lower proterozoic dome and nappe complex. descriptive text. 72 pp. copenhagen: geological survey of greenland. thomassen, b. 1991: the black angel lead-zinc mine 1973–90. rapport grønlands geologiske undersøgelse 152, 46–50. authors’ addresses p.g. & d.r., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. email: pgua@geus.dk c.a.p., department of geological sciences, university of saskatchewan, 114 science place, saskatoon, sk s7n 5e2, canada. geological survey of denmark and greenland bulletin 3, 112-151 112 fied, they offer some additional information when considered together with the sedimentological and mollusc records. conclusive remarks on the skagen well in the description given above, the faunal record is the basis for understanding the climatic changes in the skagen well, supplemented by the observation on the changes in the sediments. however, the changes found during the holocene are most likely to be connected with changes in facies, and here the changing depth is the most prominent agent, ending up with the last event represented by the depositional history of the skagen spit. based on the dating of the holocene and the late weichselian, the descriptions have been given in terms of episodes. especially the holocene strata points to a development from deeperto shallow-water facies from preboreal to subatlantic. in this development there appears to be a facies change that can be compared to the bottom communities as known from the skagerrak–kattegat region when going from the deeper-water communities of the present day, the so-called maldane-ophiura sarsi community, to the venus community of the more shallow seas. the mollusc assemblages in the skagen sequence indicate a deeper-water facies during the eemian, the weichselian and the older holocene in contrast to what hitherto was known in other parts of the danish area during the late quaternary. the skagen well has a record of the changing seas during the late quaternary, from the eemian through the weichselian (although only in parts) and the holocene. for the first time within the danish area, the full record of the marine environmental transition from the late pleistocene to the holocene can be demonstrated on the basis of molluscs. however, not all the episodes known from the skagen well can be found in marine facies of the other regions, but thanks to the new records from the north sea around the jydske rev area, a near to full holocene marine record is at hand, including part of the preboreal (petersen 1998). the environmental changes through time in the seven sectors based on the molluscan records the recorded mollusc species within each area are given in appendix 6. regarding the environmental changes through time within the danish realm, the seven sectors will be considered from the eemian, starting in the south within the classical area where forchhammer named the deposits the cyprina clay. eemian species sorted after climatic affinities the bælt sea age: eemian climatic regions: asbl class bivalvia subclass heterodonta order myoida mya truncata linnaeus 1758 total for climatic regions asb. : 1 (1.7%) climatic regions: .sb. class bivalvia subclass heterodonta order veneroida arctica islandica (linnaeus 1767) order myoida zirfaea crispata (linnaeus 1758) total for climatic regions .sb. : 2 (3.4%) climatic regions: .sbl class bivalvia subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 order pterioida geus bulletin no 3.pmd 28-06-2004, 08:45112 113 heteranomia squamula (linnaeus 1758) subclass heterodonta order veneroida mysella bidentata (montagu 1803) tellimya ferruginosa (montagu 1803) spisula elliptica (brown 1827) macoma balthica (linnaeus 1758) total for climatic regions .sbl: 6 (10.2%) climatic regions: ..bl class gastropoda subclass prosobranchia order neotaenioglossa littorina littorea (linnaeus 1758) hydrobia ulvae (pennant 1777) rissoa inconspicua alder 1844 rissoa membranacea (j. adams 1800) rissoa parva (da costa 1779) caecum glabrum (montagu 1803) bittium reticulatum (da costa 1778) aporrhais pespelicani (linnaeus 1758) order heterogastropoda triphora adversa (montagu 1803) epitonium clathrus (linnaeus 1758) order neogastropoda hinia pygmaea (lamarck 1822) hinia reticulata (linnaeus 1758) subclass heterobranchia order heterostropha odostomia scalaris macgillivray 1843 chrysallida obtusa (brown 1827) chrysallida spiralis (montagu 1803) ebala nitidissima (montagu 1803) turbonilla crenata (brown 1827) turbonilla lactea (linné 1758) subclass opisthobranchia order bullomorpha philine aperta (linnaeus 1767) order anaspidea retusa truncatula (bruguière 1792) retusa umbilicata (montagu 1803) akera bullata müller 1776 class bivalvia subclass palaeotaxodonta order nuculoida nucula nitidosa winckworth 1930 nucula sulcata (bronn 1831) subclass pteriomorpha order mytiloida modiolula phaseolina (philippi 1844) order pterioida ostrea edulis linnaeus 1758 subclass heterodonta order veneroida lepton nitidum (turton 1822) acanthocardia echinata (linnaeus 1758) parvicardium exiguum (gmelin 1791) cerastoderma edule (linnaeus 1758) mactra stultorum (linnaeus 1758) spisula subtruncata (da costa 1778) ensis ensis (linnaeus 1758) angulus tenuis (da costa 1778) scrobicularia plana (da costa 1778) abra alba (wood 1802) abra prismatica (montagu 1803) chamelea striatula (da costa 1778) timoclea ovata (pennant 1777) dosinia lincta (montagu 1803) order myoida corbula gibba (olivi 1792) barnea candida (linnaeus 1758) subclass anomalodesmata order pholadomyoida thracia phaseolina (lamarck 1818) total for climatic regions ..bl: 43 (72.9%) climatic regions: ...l class gastropoda subclass opisthobranchia order bullomorpha haminoea navicula (da costa 1778) class bivalvia subclass pteriomorpha order mytiloida mytilaster solidus (poli 1795) subclass heterodonta order veneroida lucinella divaricata (linnaeus 1758) gastrana fragilis (linnaeus 1758) abra segmentum (récluz 1843) paphia senescens (cocconi 1873) gouldia minima (montagu 1803) total for climatic regions ...l: 7 (11.9%) total for the eemian bælt sea: 59 (23.9%) fifty-nine species have been found in the bælt sea region during the eemian, seven of which are found or have been, as is the case with paphia aurea senescens, in the lusitanian region. all of the lusitanian species are found only in the eemian deposits and represent species living in the shallow-water environgeus bulletin no 3.pmd 28-06-2004, 08:45113 114 ment. mytilaster solidus is intertidal attached to rocks or algae, abra segmentum is infralittoral on sandy mud, and haminoea navicula found in zostera beds in sheltered areas. the other three living species lucinella divaricata, gastrana fragilis and gouldia minima are found from or just below the tidal zone and further out at different depths. paphia aurea senescens may as well be regarded as a shallow-water species, considering the distribution of the other tapes species. in this way the overall climatic characterising species of the eemian in the bælt sea area are connected with the shallow-water environment. all the boreo-lusitanian species are known from the recent danish fauna. this is by far the largest group of molluscs, with 43 species forming 72.9% of the eemian bælt sea fauna. 33 species can be found in the tidal to shallow-water environment, while 11 species – epitonium clathrus, ebala nitidissima, turbonilla crenata, retusa umbilicata, akera bullata, nucula nitidosa, nucula sulcata, lepton nitidum, acanthocardia echinata, phaxas pellucidus and timoclea ovata – find their minimum depth, defined by acanthocardia echinata, phaxas pellucidus and timoclea ovata, at 4 m, and epitonium clathrus at 5 m. so rather considering the maximum depth indicated by some of the shallow-water species, there must be two faunas, of which one is shallow out to a depth of a few metres and another for deeper water. there are six species with a rather broad range from the subarctic to the lusitanian regions. five of these species can be found in shallow water, including the intertidal zone, except spisula elliptica, which occurs only at greater depth. the eulittoral species mytilus edulis has given name to the mytilus beds found in the lower part of the marine eemian deposits in the bælt sea region, characterising the littoral deposits. only two species, arctica islandica and zirfaea crispata from the bælt sea region, have a subarctic–boreal distribution. the overall characteristic species for the eemian bælt sea deposits – arctica islandica – can be found from the intertidal zone to great depth, but in the inner recent danish waters often at depths from 10 to 60 m. in the eemian this species characterises the clay deposited during the deeper-water facies. the other mainly boreal species zirfaea crispata has been found only in the upper part of the tapes sand at stensigmose (madsen et al. 1908, p. 176) which fits well with the depth interval of this species from low tide to 7 m. finally mya truncata covers the arctic, subarctic, boreal and lusitanian regions down to the bay of biscay and can be found from the intertidal zone down to 75 m, in danish waters often between 10 and 20 m. taken together, all the information from the abovementioned climatic groups indicates that the bælt sea deposits are represented by two facies. one in littoral/ infralittoral water not deeper than maximum 10 m, and one at depths of more than 5 m, and it is seen that all of the characteristic eemian species within the bælt sea area are lusitanian species connected with the shallow-water facies. fifteen mollusc species of the eemian bælt sea fauna show by the region of lowest mean salinity they inhabit that the salinity of the bælt sea area must have been higher than present-day waters by up to 30–33‰ (sorgenfrei 1958, table 11). here listed as mentioned by nordmann (1928, pp. 79–81): 33‰ circe minima (gouldia minima), 31‰ dosinia lincta, 33‰ lepton nitidum, 33‰ mactra stultorum, 30‰ montacuta ferruginosa (tellimya ferruginosa), 30‰ mytilus phaseolinus (modiolula phaseolina), 33‰ nucula sulcata, 31‰ syndosmya prismatica (abra prismatica), 30‰ venus gallina (chamelea striatula), 30‰ parthenia interstincta (chrysallida spiralis), 32‰ rissoa parva, 30‰ (epitonium clathrus), 33‰ turbonilla lactea, and 30‰ turbonilla rufa (turbonilla crenata). in the recent bottom communities out in the bælt sea area, macoma balthica is the overall characteristic mollusc occurring in all the samples from this region, in some cases being the only mollusc and in some cases together with others such as mytilus edulis, cerastoderma edule and scrobicularia plana in shallow water and in places with a vegetation of gastropods like littorina littorea, littorina tenebrosa, rissoa membranacea and rissoa inconspicua (petersen 1913). the littoral elements with mytilus edulis are well documented in the eemian deposits by the so-called mytilus beds, together with the infaunal paphia aurea senescens found in situ within the strata around the mytilus beds or even down in the freshwater layers. however, macoma balthica is only found in the tapes sands at stensigmose in southern jutland (fig. 1) and not at the many other bælt sea localities (nordmann 1908, 1928). in the deeper-water community of today the macoma community is replaced by the abra alba or astarte communities (petersen 1913). here again the eemian deposits differ on the leading species in that fig. 99. the voderup klint section on the island of ærø with dislocated marine clays of eemian age. geus bulletin no 3.pmd 28-06-2004, 08:45114 115 geus bulletin no 3.pmd 28-06-2004, 08:45115 116 fig. 100. one of the cyprina clay outcrops along the voderup klint profile on the island of ærø. abra alba is recorded only from stensigmose and the astartidae tridonta borealis, tridonta elliptica and tridonta montagui have not been found at all in the bælt sea region or further to the east within the baltic. the characteristic species of the eemian deeperwater deposits, the cyprina clay (with an abundant number of arctica islandica), is also present in the recent deeper water of the bælt sea, but petersen (1913, p. 4) avoided using this bivalve as one of his characterising species in his ‘evaluation of the sea’ because of its uneven distribution, which is certainly not the case considering the present outcrops of the cyprina clay along the shores of the bælt sea region (figs 99, 100). the great difference between the eemian bælt sea fauna and the recent one which has been outlined above can be considered together with the concluding remarks of nordmann (1908, pp. 113 and 148) on cyprina islandica, stating that this species cannot be regarded as a characterising fossil species of the eemian deposits, but more probably should be seen as a relict in the inner part of the eemian waters from the sea, of a more arctic–boreal nature, so to say forming a parallel to the occurrences of the astarte species in the present bælt sea and baltic regions. the baltic age: eemian climatic regions: asb. class bivalvia subclass heterodonta order veneroida macoma calcarea (gmelin 1791) total for climatic regions asb. : 1 (5.3%) climatic regions: asbl class bivalvia subclass heterodonta order myoida mya truncata linnaeus 1758 total for climatic regions asbl: 1 (5.3%) climatic regions: .sb. class bivalvia subclass heterodonta order veneroida arctica islandica (linnaeus 1767) total for climatic regions .sb. : 1 (5.3%) climatic regions: .sbl class bivalvia subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 modiolus modiolus (linnaeus 1758) subclass heterodonta order veneroida mysella bidentata (montagu 1803) macoma balthica (linnaeus 1758) total for climatic regions .sbl: 4 (21.1%) climatic regions: ..bl class gastropoda subclass prosobranchia order neotaenioglossa littorina littorea (linnaeus 1758) hydrobia ulvae (pennant 1777) turritella communis risso 1826 aporrhais pespelicani (linnaeus 1758) lunatia alderi (forbes 1838) order neogastropoda hinia reticulata (linnaeus 1758) class bivalvia subclass palaeotaxodonta order nuculoida nucula nitidosa winckworth 1930 nucula nucleus (linnaeus 1767) subclass heterodonta order veneroida cerastoderma edule (linnaeus 1758) spisula subtruncata (da costa 1778) scrobicularia plana (da costa 1778) order myoida corbula gibba (olivi 1792) total for climatic regions ..bl: 12 (63.2%) total for the eemian baltic: 19 (7.7%) the baltic sea fauna comprises only 19 species. however, five are new here compared to the eemian known from the bælt sea area: three from the boreo-lusitanian climatic region, i.e. turritella communis, lunatia alderi and nucula nucleus; one from the subarctic– lusitanian region, i.e. modiolus modiolus; and one with an arctic–boreal distribution, i.e. macoma calcarea. according to the living depth of turritella communis geus bulletin no 3.pmd 28-06-2004, 08:45116 117 geus bulletin no 3.pmd 28-06-2004, 08:45117 118 and lunatia alderi, this should be an indication of environment deeper than 10 m. but the eulittoral mytilus edulis and cerastoderma edule are nevertheless recorded from the baltic sea area, although these species do not come from layers in which they occur in great quantities in the bælt sea area (the mytilus beds). from the climatic affinities of the molluscan finds in the baltic sea area, it appears that no pure lusitanian species are found. this is the biggest contrast to the bælt sea region. it has also been the basis for keeping these finds apart, as done by ødum (1933). however, considering the new element in this fauna, turritella communis, compared to the bælt sea fauna, it has been argued by way of correlation on the basis of the foraminifera that these deposits can be regarded as eemian and that the facies belongs in deeper water than known from the bælt sea area (petersen & konradi 1974). such deposits occur, besides that at strandegaard dyrehave, also at some places on møn characterised by the occurrences of turritella communis (berthelsen et al. 1977). in the region of lowest mean salinity inhabited (sorgenfrei 1958, table 11), for 5 species from the baltic during the eemian, i.e. modiolus modiolus, turritella communis, nucula nitidosa, nucula nucleus, and spisula subtruncata, a salinity between 20‰ and 25‰ can be shown, which is above the present conditions in the baltic. however, this might be caused by a higher tide in inner danish waters during the eemian, in this way giving the same situation as known during the atlantic, when the salinity was higher. unlike the situation in the bælt sea, the present-day characterising species macoma balthica is recorded from the localities on sjælland, strandegaards dyrehave and møn. the drop in number of species among bivalves and gastropods from the bælt sea to the baltic is from 45 to 11 species respectively in the present-day fauna (sorgenfrei 1958), which is of the same order of magnitude as seen in the fossil fauna from the bælt sea to the baltic during the eemian. this means that in some waywe have to do with the same basin structure/hydrographic situation, although the salinity was higher during the eemian in the innermost danish waters than at present. the kattegat age: eemian climatic regions: asbl class bivalvia subclass heterodonta order myoida hiatella arctica (linnaeus 1758) total for climatic regions asbl: 1 (3.1%) climatic regions: .sbl class gastropoda subclass prosobranchia order archaeogastropoda scissurella crispata fleming 1828 order neogastropoda buccinum undatum linnaeus 1758 class bivalvia subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 total for climatic regions .sbl: 3 (9.4%) climatic regions: ..bl class gastropoda subclass prosobranchia order neotaenioglossa littorina littorea (linnaeus 1758) hydrobia ulvae (pennant 1777) rissoa inconspicua alder 1844 rissoa violacea desmarest 1814 bittium reticulatum (da costa 1778) turritella communis risso 1826 lunatia alderi (forbes 1838) order neogastropoda hinia pygmaea (lamarck 1822) hinia reticulata (linnaeus 1758) subclass heterobranchia order heterostropha chrysallida spiralis (montagu 1803) turbonilla lactea (linné 1758) class bivalvia subclass pteriomorpha order pterioida ostrea edulis linnaeus 1758 subclass heterodonta order veneroida acanthocardia echinata (linnaeus 1758) parvicardium exiguum (gmelin 1791) cerastoderma edule (linnaeus 1758) geus bulletin no 3.pmd 28-06-2004, 08:45118 119 laevicardium crassum (gmelin 1791) spisula solida (linnaeus 1758) spisula subtruncata (da costa 1778) scrobicularia plana (da costa 1778) abra prismatica (montagu 1803) paphia aurea (gmelin 1791) timoclea ovata (pennant 1777) order myoida corbula gibba (olivi 1792) total for climatic regions ..bl: 23 (71.9%) climatic regions: ...l class gastropoda subclass opisthobranchia order bullomorpha haminoea navicula (da costa 1778) class bivalvia subclass pteriomorpha order mytiloida mytilaster lineatus (gmelin 1791) subclass heterodonta order veneroida lucinella divaricata (linnaeus 1758) gastrana fragilis (linnaeus 1758) paphia senescens (cocconi 1873) total for climatic regions ...l: 5 (15.6%) the eemian kattegat: 32 (13.0%) from the region of the kattegat sensu lato area, including the bordering landmasses with fjords, sounds and the northern part of the lillebælt, storebælt, and øresund, three kinds of localities have been met with. as the information obtained from this area is based on a very different kind of material, the localities will be treated apart. the first locality to be considered is ejby bro in northern sjælland on the isefjord. here erik rasmussen (in madsen 1968) described an in situ marine deposit. out of 15 mollusc species, two are lusitanian, viz. lucinella divaricata and paphia aurea senescens, both of which are characteristic fossils of the eemian. furthermore, out of ten boreo-lusitanian species, seven are not recorded from the recent isefjord, viz. rissoa violacea, lunatia alderi, hinia reticulata (the british form (fretter & graham 1984, p. 495)), ostrea edulis, laevicardium crassum, paphia aurea, and spisula solida. among these are also the less tolerant species regarding the region of lowest mean salinity inhabited, which is five species between 28‰ and 33‰. the present-day figure of salinity for the isefjord is between 18 and 20‰. although we have to do with a saalien glacial topography, we may explain the higher salinity as a result of a higher tidal amplitude during the eemian. only two species occurring in the ejby bro locality extend into the subarctic: buccinum undatum and mytilus edulis, besides one – hiatella arctica – having a wide range. all are present also in the recent isefjord waters. this fauna points to a more oceanic environment than today and with a higher temperature; furthermore, the finds point to shallow-water or even beach deposits with some tidal influence (rasmussen in madsen 1968). the second kind of locality includes redeposited sediments: either floes in the glacial deposits like the stautrup locality at aarhus or fluvioglacial deposits as at høng in western sjælland (nordmann 1928, pp. 64– 65; ødum 1933; sorgenfrei 1945). both of these localities carried some of the characteristic molluscs of the eemian deposits. from høng the following have been recorded: lucinella divaricata and paphia aurea senescens together with 11 species known from the boreo-lusitanian zone and mytilus edulis known from the subarctic to the lusitanian regions. only four species were also found at the ejby bro locality, and the ‘new’ ones (nine species) point to deeper water with such species as acanthocardia echinata, corbula gibba and turritella communis. the stautrup material records new species to the kattegat region such as rissoa inconspicua, mytilaster lineatus, gastrana fragilis, abra prismatica, chrysallida spiralis, and haminoea navicula, all except mytilaster lineatus and gastrana fragilis being boreo-lusitanian, while the two bivalves are limited to the lusitanian and characteristic of the eemian – together with paphia aurea senescens, which is also present in the stautrup floe, as discussed by sorgenfrei (1945). from the material hitherto discussed it appears that the eemian sea deposits were known from the kattegat region both in a shallow-water facies (ejby bro) and a deeper-water facies (høng), the latter only from redeposited material. therefore, it is of great importance from a palaeogeographical point of view that information now has been obtained from borings in the central part of the kattegat, on the island of anholt (lykke-andersen et al. 1993). foraminifera from this well (seidenkrantz 1993) revealed a marine upper saalian and eemian sequence. the macrofossils from the well were kindly placed at my disposal. it appears that a temperate fauna with species such as turritella geus bulletin no 3.pmd 28-06-2004, 08:45119 120 communis, aclis minor, hiatella arctica, nuculana minuta, and scissurella crispata is resting on an arctic deposit with portlandia arctica. the information that we do have is that a turritella facies within the kattegat region in the eemian found in an in situ position sustains the view of a continuation towards the north not only of the shallow-water deposits, but also of the deeper-water environment during the eemian. considering the bottom communities of the present-day kattegat as revealed by petersen (1913), turritella communis is found in different associations at depths from 12–19 to 35 m on sand, fine sand and clay. the species from the eemian of the kattegat associated with turritella communis in the kattegat of today are hiatella arctica, nuculana minuta, acanthocardia echinata, corbula gibba, and hinia pygmaea. considering the eemian faunas demonstrated from the kattegat region, it appears that here the lowest mean salinity inhabited by the species in question is within the present-day salinities reached for this area, at about 33‰, although the salinities for the present fjords bordering the kattegat have a lower salinity, as mentioned in the case of the isefjord with the ejby bro locality. the north sea age: eemian climatic regions: asb. class bivalvia subclass heterodonta order veneroida tridonta elliptica (brown 1827) total for climatic regions asb. : 1 (1.1%) climatic regions: asbl class gastropoda subclass prosobranchia order neotaenioglossa littorina saxatilis (olivi 1792) class bivalvia subclass heterodonta order veneroida thyasira flexuosa (montagu 1803) tridonta montagui (dillwyn 1817) order myoida mya truncata linnaeus 1758 hiatella arctica (linnaeus 1758) total for climatic regions asbl: 5 (5.5%) climatic regions: .sb. class bivalvia subclass heterodonta order veneroida arctica islandica (linnaeus 1767) total for climatic regions .sb. : 1 (1.1%) climatic regions: .sbl class gastropoda subclass prosobranchia order neotaenioglossa littorina obtusata (linnaeus 1758) lacuna vincta (montagu 1803) order neogastropoda buccinum undatum linnaeus 1758 class bivalvia subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 modiolus modiolus (linnaeus 1758) order pterioida heteranomia squamula (linnaeus 1758) subclass heterodonta order veneroida mysella bidentata (montagu 1803) tellimya ferruginosa (montagu 1803) parvicardium ovale (sowerby 1840) macoma balthica (linnaeus 1758) total for climatic regions .sbl: 10 (11.0%) climatic regions: ..bl class gastropoda subclass prosobranchia order archaeogastropoda gibbula cineraria (linnaeus 1758) order neotaenioglossa littorina littorea (linnaeus 1758) lacuna parva (montagu 1803) hydrobia ulvae (pennant 1777) onoba vitrea (montagu 1803) rissoa albella lovén 1846 rissoa inconspicua alder 1844 rissoa membranacea (j. adams 1800) rissoa parva (da costa 1779) rissoa violacea desmarest 1814 caecum glabrum (montagu 1803) bittium reticulatum (da costa 1778) turritella communis risso 1826 aporrhais pespelicani (linnaeus 1758) lunatia alderi (forbes 1838) order heterogastropoda geus bulletin no 3.pmd 28-06-2004, 08:45120 121 triphora adversa (montagu 1803) cerithiopsis tubercularis (montagu 1803) epitonium clathrus (linnaeus 1758) order neogastropoda hinia pygmaea (lamarck 1822) hinia reticulata (linnaeus 1758) cytharella coarctata (forbes 1840) subclass heterobranchia order heterostropha brachystomia eulimoides hanley 1844 odostomia scalaris macgillivray 1843 chrysallida obtusa (brown 1827) chrysallida spiralis (montagu 1803) ebala nitidissima (montagu 1803) odostomia albella lovén 1846 turbonilla crenata (brown 1827) turbonilla lactea (linné 1758) subclass opisthobranchia order bullomorpha acteon tornatilis (linnaeus 1758) philine aperta (linnaeus 1767) order anaspidea retusa truncatula (bruguière 1792) akera bullata müller 1776 class bivalvia subclass palaeotaxodonta order nuculoida nucula nitidosa winckworth 1930 nucula sulcata (bronn 1831) subclass pteriomorpha order mytiloida modiolula phaseolina (philippi 1844) modiolaria tumida (hanley 1843) order pterioida aequipecten opercularis (linnaeus 1758) chlamys varia (linnaeus 1758) ostrea edulis linnaeus 1758 subclass heterodonta order veneroida lepton nitidum (turton 1822) acanthocardia echinata (linnaeus 1758) parvicardium exiguum (gmelin 1791) parvicardium scabrum (philippi 1844) cerastoderma edule (linnaeus 1758) spisula subtruncata (da costa 1778) ensis ensis (linnaeus 1758) phaxas pellucidus (pennant 1777) angulus tenuis (da costa 1778) tellina donacina linnaeus 1758 fabulina fabula (gmelin 1791) donax vittatus (da costa 1778) scrobicularia plana (da costa 1778) abra alba (wood 1802) abra prismatica (montagu 1803) chamelea striatula (da costa 1778) tapes decussatus (linnaeus 1758) timoclea ovata (pennant 1777) venerupis pullastra (montagu 1803) dosinia lincta (montagu 1803) mysia undata (pennant 1777) order myoida corbula gibba (olivi 1792) saxicavella jeffreysi winckworth 1930 barnea candida (linnaeus 1758) subclass anomalodesmata order pholadomyoida thracia phaseolina (lamarck 1818) thracia villosiuscula (macgillivray 1827) total for climatic regions ..bl: 66 (72.5%) climatic regions: ...l class gastropoda subclass opisthobranchia order bullomorpha haminoea navicula (da costa 1778) class bivalvia subclass pteriomorpha order mytiloida mytilaster lineatus (gmelin 1791) subclass heterodonta order veneroida lucinella divaricata (linnaeus 1758) plagiocardium papillosum poli 1795 gastrana fragilis (linnaeus 1758) abra segmentum (récluz 1843) paphia senescens (cocconi 1873) gouldia minima (montagu 1803) total for climatic regions ...l: 8 (8.8%) the eemian north sea: 91 (36.8%) regarding the eemian deposits of south-western jylland several localities are included: tønder and surroundings, forballum, farup, ydre bjergum, mandø hølade (many places) and inder bjergum. furthermore, also molluscs of eemian age are recorded from the map sheet blaavands huk (fig. 1). in all, 91 molluscan species have been recorded, 53 bivalves and 38 gastropods. this high number of species comprises 32 species new to the eemian compared to the bælt sea deposits. geus bulletin no 3.pmd 28-06-2004, 08:45121 122 considering the new elements from the climatic point of view, two are lusitanian species, plagiocardium papillosum and mytilaster lineatus, which are regarded as part of the characteristic species of the eemian fauna (nordmann 1928). by far the largest group of new species in the north sea deposits are the boreo-lusitanian. only one – tridonta elliptica – does not extend into the lusitanian region. this is close to the situation found in the bælt sea region, where only two species; arctica islandica and zirfaea crispata, do not reach the lusitanian region. arctica islandica – so common in the bælt sea deposits – has been recorded only in a single find of a juvenile specimen at mandø hølade in south-western jylland. the other species, zirfaea crispata, has not been demonstrated at all in the danish north sea eemian deposits. the boreo-lusitanian species mactra stultorum has a high salinity requirement, occurring in the bælt sea region but not in the danish north sea region during the eemian. this has been used as an argument against uniting in time the deposits found on the western and eastern sides of southern jylland. “on a voulu y voir une preuve que ces deux bassins de mer n’out en réalité rien eu à faire l’un avec l’autre, en sorte qu’ils pourraient très bien être d’àges fort différents” (nordmann 1928, p. 63). when the 15 mollusc species of the eemian bælt sea fauna mentioned earlier with a salinity requirement between 30–33‰ are remembered, the former focus on mactra stultorum is of less significance. while nordmann states that the fauna on both sides of the jylland peninsula can be regarded as one, he is right from the point of view of climatic conditions, as demonstrated above. however, he is also arguing for a connection between the north sea and the bælt sea (nordmann 1928, p. 63): “les passes entre les parties orientale et occidentale de la mer eemienne, ce qui, autrement parlant, signifie les vallées et les plaines entre les collines insulaires qui sont aujourd’hui occupées par les plaines de landes du slesvig et du holstein, sont sans doute très etroites”. however, it is the author’s opinion that the difference between the southern danish localities to the east (the bælt sea) and to the west (the north sea) can be regarded as differences in facies that are also found in present-day danish waters. here the occurrences of donax vittatus in the eemian north sea deposits, but not in the bælt sea deposits, should be considered. donax vittatus is a characteristic species of the high-energy coastal environment, where it is found all along the present-day west coast, but not in the inner danish waters from the northernmost part of the east coast of jylland. however, many of the other molluscs might have made their way to the bælt sea and the baltic region if there has been only the slightest passage over southern jylland. in the case of a passage, the situation can be looked upon as a parallel to the present-day marine colonisation of the limfjord after the breaking through by the north sea at the spit at agger in 1825. at the end of the 19th century a rich marine fauna could be recorded in the limfjord (collin 1884; petersen 1888). the question of connection between the north sea and the bælt sea over the southern part of the peninsula of jylland can be considered, based on the mollusc assemblages. it appears that the difference in faunal composition during the eemian in both seas was very similar to the difference in faunas during the holocene, although with lusitanian shallow-water species occurring during the eemian. therefore it must be concluded that the land–sea configuration must be very much the same during the two periods so jylland was also a peninsula during the eemian. vendsyssel age: eemian climatic regions: asb. class gastropoda subclass prosobranchia order neogastropoda oenopota incisula (verrill 1882) oenopota violacea (mighels & adams 1842) bela exarata g.o. sars 1818 class bivalvia subclass palaeotaxodonta order nuculoida nuculana pernula (müller 1776) yoldiella lenticula (möller 1842) subclass pteriomorpha order mytiloida musculus niger (gray 1824) subclass heterodonta order veneroida clinocardium ciliatum (fabricius 1780) serripes groenlandicus (bruguière 1798) macoma calcarea (gmelin 1791) total for climatic regions asb. : 9 (16.4%) geus bulletin no 3.pmd 28-06-2004, 08:45122 123 climatic regions: asbl class gastropoda subclass prosobranchia order neogastropoda oenopota trevelliana (turton 1834) subclass opisthobranchia order thecosomata limacina retroversa (fleming 1823) class bivalvia subclass palaeotaxodonta order nuculoida nuculoma tenuis (montagu 1808) yoldiella frigida (torell 1859) subclass heterodonta order veneroida leptaxinus ferruginosus (forbes 1844) order myoida mya truncata linnaeus 1758 hiatella arctica (linnaeus 1758) total for climatic regions asbl: 7 (12.7%) climatic regions: .sb. class bivalvia subclass palaeotaxodonta order nuculoida nuculana minuta (müller 1776) total for climatic regions .sb. : 1 (1.8%) climatic regions: .sbl class gastropoda subclass prosobranchia order neotaenioglossa lacuna vincta (montagu 1803) order neogastropoda buccinum undatum linnaeus 1758 class bivalvia subclass palaeotaxodonta order nuculoida yoldiella lucida (lovén 1846) subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 subclass heterodonta order veneroida mysella bidentata (montagu 1803) parvicardium ovale (sowerby 1840) total for climatic regions .sbl: 6 (10.9%) climatic regions: ..b. class gastropoda subclass heterobranchia order heterostropha chrysallida eximia (jeffreys 1849) total for climatic regions ..b. : 1 (1.8%) climatic regions: ..bl class gastropoda subclass prosobranchia order neotaenioglossa alvania abyssicola (forbes 1850) rissoa parva (da costa 1779) bittium reticulatum (da costa 1778) turritella communis risso 1826 aporrhais pespelicani (linnaeus 1758) lunatia alderi (forbes 1838) order heterogastropoda vitreolina philippii (rayneval & ponzi1854) order neogastropoda hinia incrassata (ström 1768) hinia reticulata (linnaeus 1758) mangelia brachystoma (philippi 1844) raphitoma linearis (montagu 1803) subclass heterobranchia order heterostropha eulimella scillae (scacchi 1835) ondina divisa (j. adams 1797) odostomia turrita hanley 1844 subclass opisthobranchia order bullomorpha acteon tornatilis (linnaeus 1758) order anaspidea retusa umbilicata (montagu 1803) class bivalvia subclass palaeotaxodonta order nuculoida nucula nucleus (linnaeus 1767) nucula sulcata (bronn 1831) yoldiella philippiana (nyst 1845) subclass pteriomorpha order pterioida pseudamussiumseptemradiatum (müller 1776) similipecten similis (laskey 1811) subclass heterodonta order veneroida acanthocardia echinata (linnaeus 1758) parvicardium minimum (philippi 1836) ensis ensis (linnaeus 1758) phaxas pellucidus (pennant 1777) abra alba (wood 1802) abra nitida (müller 1776) abra prismatica (montagu 1803) geus bulletin no 3.pmd 28-06-2004, 08:46123 124 kelliella miliaris (philippi 1844) chamelea striatula (da costa 1778) order myoida corbula gibba (olivi 1792) total for climatic regions ..bl: 31 (56.4%) the eemian vendsyssel: 55 (22.3%) in vendsyssel the mollusc faunas in the skærumhede sequence have been studied by nordmann (jessen et al. 1910) and petersen (bahnson et al. 1974). the conclusion reached in the latter study on molluscs points out that the difference between the boreo-lusitanian community – the turritella terebra zone – in the boring and the typical eemian community in southern denmark is a difference in facies. this statement will be discussed now on the basis of all the mollusc found and listed according to their climatic regions. among the 55 species of molluscs recorded from the eemian in the vendsyssel region, no lusitanian species occur, but boreo-lusitanian species count for more than half of the assemblages (31 species, 56.4%). of these, 25 are mentioned in the list for region of lowest mean salinity inhabited built on information from the danish waters in the transition area between the north sea (at esbjerg) and the baltic (gulf of bothnia) (sorgenfrei 1958). nearly 4/5 of this number have their region of lowest mean salinity between 24‰ and 34‰, which is the minimum, and mean salinities at the passage belt between the north sea – skagerrak and e and nw kattegat. this situation for an eemian assemblage indicates a high degree of similarity with the present-day environment in this area. within the boreo-lusitanian group of molluscs there is a clear dominance of species belonging to the deeperwater environment at around the 100 m depth. however, there are some few species which belong at a depth of less than 20–30 m. such species are rissoa parva, hinia reticulata and retusa umbilicata, which within their depth range live in great abundance, which is not the case in the skærumhede sequences. therefore they can be regarded as allochthonous or as stray finds outside their environment. nordmann (in jessen et al. 1910) mentioned that rissoa parva and hinia reticulata were redeposited. he also mentioned bittium reticulatum, which according to the literature has some records from the deeper water (out to 250 m deep). taking the whole group of boreo-lusitanian species, there are many which occur in the tidal and shallow-water environments but have a wide range of depth. rather few are those which are mainly connected with the deeper water – here depths of more than 10 m – viz.: turritella communis, lunatia alderi, mangelia brachystoma, raphitoma linearis, eulimella scillae, nucula sulcata, pseudamussium septemradiatum and kelliella miliaris. the last one strengthens the depth indications to be more than 100 m, since the species is recently known only from the deeper part of the skagerrak. chrysallida eximia, which is found at depths from 20–1000 m is here regarded as a boreal species, although it extends into the northernmost part of the lusitanian zone, but there in deeper water. among the species known from the subarctic to the lusitanian regions, the finds of mytilus edulis (jessen et al. 1910; bahnson et al. 1974) can be regarded from the same point of view as mentioned in connection with rissoa parva and hinia reticulata. the few finds of mytilus specimens clearly indicate that this normally gregariously living species is found outside its living zone and can be regarded also as stray finds. the other finds of species occurring in the subarctic to the lusitanian zone do not oppose the view of a deeper-water environment. taking the species together which do not occur in the lusitanian zone but extend into the arctic or subarctic (nuculana minuta), there are two species clinocardium ciliatum and serripes groenlandicus, which do pose a problem regarding their climatic indications within the skærumhede sequence. these two species are found only in the high boreal zone (norway north of lofoten). nordmann (jessen et al. 1910, pp. 127– 128) mentioned the occurrences of these species off iceland together with other species with which they occur in the skærumhede sequence. this picture has been to some degree changed by the publication of parts of the zoology of iceland, so it appears that turritella communis is no longer regarded as being part of the icelandic fauna, while eulimella scillae (only empty shells) might be added. lunatia alderi and raphitoma linearis mentioned by nordmann are confirmed (thorson 1941), and so is acteon tornatilis (lemche 1938), in the new literature. however, in this way we still face the question of the climatic conditions indicated by the molluscan assemblage. however, as seen from the latest skærumhede boring (bahnson et al. 1974, fig. 7), there is a clear transition zone between the turritella communis zone and the establishment of the turritella erosa zone. in this zone, the so-called abra nitida zone, we still find the geus bulletin no 3.pmd 28-06-2004, 08:46124 125 mixing of species with different climatic affinities, just as we do not have a sharp border zone to tell where we actually leave the eemian and pass into the weichselian. but we do have a well-defined bottom community for what we must call the eemian from the vendsyssel region, and that is the turritella communis community. this community we find in deeper water, also in the recent danish waters. oenopota incisula is one of the species also occurring in the transition zone between the two turritella communities, the boreal–lusitanian with t. communis, and the arctic with t. erosa. oenopota incisula has been found in both borings (jessen et al. 1910; bahnson et al. 1974, fig. 7), but the species has still not been recorded from the european coast of the north atlantic in recent time. in this way oenopota incisula becomes one of the few species extinct in our part of the world since the early/middle weichselian. finally the six species with a wide range within the climatic regions are all found in deeper water. however, here yoldiella frigida occurs in deeper water in the southern part of its range. limacina retroversa is a pelagic species and has been recorded in the present day to penetrate into the kattegat – bælt sea regions. skagen age: eemian climatic regions: asb. class bivalvia subclass palaeotaxodonta order nuculoida nuculana pernula (müller 1776) total for climatic regions asb. : 1 (7.1%) climatic regions: asbl class gastropoda subclass opisthobranchia order thecosomata limacina retroversa (fleming 1823) class scaphopoda siphonodentaliumlobatum (sowerby 1860) class bivalvia subclass palaeotaxodonta order nuculoida yoldiella frigida (torell 1859) subclass heterodonta order myoida hiatella arctica (linnaeus 1758) total for climatic regions asbl: 4 (28.6%) climatic regions: .sbl class scaphopoda antalis entalis (linnaeus 1758) class bivalvia subclass pteriomorpha order pterioida delectopecten vitreus (gmelin 1791) total for climatic regions .sbl: 2 (14.3%) climatic regions: ..bl class gastropoda subclass prosobranchia order neotaenioglossa lunatia alderi (forbes 1838) subclass opisthobranchia order bullomorpha philine catena (montagu 1803) class scaphopoda cadulus subfusiforme (m. sars 1865) entalina tetragona (brocchi 1814) class bivalvia subclass heterodonta order veneroida kelliella miliaris (philippi 1844) subclass anomalodesmata order pholadomyoida cochlodesma praetenue (pulteney 1799) total for climatic regions ..bl: 6 (42.9%) climatic regions: ...l class scaphopoda dentalium vulgare da costa 1778 total for climatic regions ...l: 1 (7.1%) the eemian skagen: 14 (5.7%) in the eemian part of the skagen well, one species, dentalium vulgare, mainly occurs within the lusitanian region and is not recorded from recent danish waters, although it is found in the southern part of the north sea. among the six boreo-lusitanian mollusc species, cochlodesma praetenue is rare in danish waters, while all the others as far as the information on habitats goes are connected with the deeper-water environment. this is also true for the species including the subarctic region, viz.: antalis entalis and delectopecten vitreus, where the latter in general has a depth range from 30– 600 m, but in skagerrak is found between 400 and 600 m. also antalis entalis has a wide range of depth, geus bulletin no 3.pmd 28-06-2004, 08:46125 126 but within danish water it is recorded only from 20–400 m. the four species: siphonodentalium lobatum, limacina retroversa, yoldiella frigida and hiatella arctica with a wide geographical distribution from the arctic to the lusitanian also have a wide range of depth. also nuculana pernula, here listed from the arctic to the boreal, could be considered to have a wide range like the other species mentioned above, however, in sw europe it is found only at depths greater than 400 m. in recent danish waters it occurs at depths from 20 to 200 m. seen together with the other species in this region during the eemian with a more southern affinity, nuculana pernula shows accordance, considering that the species in the arctic is mostly littoral. comparing the assemblages from the faunal elements from skærumhede, it appears that the skagen well depicts a deeper-water community without any influence from more shallow-water facies. the occurrence of three of the four scaphopoda cadulus subfusiforme, entalina tetragona and antalis entalis speaks in favour of an environment which is likely to be found in the deeper skagerrak, such as the amphilepis norvegica/pecten vitreus community with entalina tetragonaand of which one of the other molluscs, delectopecten vitreus, is considered a characteristic species. the demonstrated eemian assemblages in most of the danish regions show differences in climatic affinities – more lusitanian in the southern part – which, however, can be explained through an analysis of the faunas in their relation to depth – shallow-water lusitanian species in the south – and in some parts in relation to the community. the community concept which was developed for the recent danish waters does not find its equivalent in the eemian bælt sea and baltic regions, only partly in the kattegat region, but has a good correlation with the occurrences of the faunal assemblages in the regions to the north in jylland – the vendsyssel and skagen regions – only in these two regions does the succession of strata allow us to follow the development into the weichselian cooler/arctic molluscan fauna, although we find some deposits from the kattegat region (holmstrup and holbæk sites) which can be correlated to the early/middle weichselian in the vendsyssel and skagen region. early/middle weichselian species sorted after climatic affinities the kattegat age: early/middle weichselian climatic regions: a... class bivalvia subclass palaeotaxodonta order nuculoida portlandia arctica (gray 1824) total for climatic regions a... : 1 (14.3%) climatic regions: as.. class gastropoda subclass opisthobranchia order bullomorpha cylichna occulta (mighels 1841) total for climatic regions as.. : 1 (14.3%) climatic regions: asb. class bivalvia subclass palaeotaxodonta order nuculoida nuculana pernula (müller 1776) subclass heterodonta order veneroida macoma calcarea (gmelin 1791) total for climatic regions asb. : 2 (28.6%) climatic regions: asbl class bivalvia subclass palaeotaxodonta order nuculoida nuculoma tenuis (montagu 1808) subclass heterodonta order myoida mya truncata linnaeus 1758 hiatella arctica (linnaeus 1758) total for climatic regions asbl: 3 (42.9%) the early/middle weichselian kattegat: 7 (2.8%) in the kattegat region, two localities (holbæk and holmstrup, fig. 1) have mollusc faunas of early/middle weichselian age, which could be correlated to the older yoldia clay deposits found in the vendsyssel and skagen regions. nordmann (ødum 1933) described the mollusc fauna in the borings at holbæk to represent part of the portgeus bulletin no 3.pmd 28-06-2004, 08:46126 127 landia arctica zone in the skærumhede sequence (jessen et al. 1910). later, petersen & buch (1974) referred the outcrops at holmstrup with marine clay, characterised by macoma calcarea, to the weichselian part of the sequence close to the macoma calcarea zone (bahnson et al. 1974) in the new well at skærumhede in the vendsyssel region. the later correlation was mainly based on the foraminiferal studies (buch in petersen & buch 1974). also aminostratigraphic investigations have to some extent sustained this correlation (miller & mangerud 1985, p. 261). the mollusc faunas from these two localities point to an arctic environment as seen from the climatic indications, in that all seven species are found in the arctic and only three of them with a wide range: hiatella arctica, mya truncata and nuculoma tenuis. taken into account that these deposits show an arctic affinity, the interpretation of depth ranges of the species found point to the more shallow-water environment. this is also true for species such as nuculoma tenuis, which is recorded from offshore down to 300 m deep, but in the arctic is more littoral. macoma calcarea is found intertidal to several hundred metres, but only in deeper water in the southern part of the range for this species. in the arctic it is the characterising mollusc in shallow water: the arctic macoma calcarea community. furthermore, nuculana pernula, as mentioned earlier, is littoral in the arctic. the conclusion to draw from these finds in the kattegat region is that we have a part – the more shallow water – of the arctic zones recorded from the vendsyssel region represented within the kattegat region – southern part. vendsyssel age: early/middle weichselian climatic regions: a... class bivalvia subclass palaeotaxodonta order nuculoida portlandia arctica (gray 1824) total for climatic regions a... : 1 (2.8%) climatic regions: as.. class gastropoda subclass prosobranchia order neotaenioglossa alvania cruenta odhner 1915 turritella erosa couthouy 1838 subclass opisthobranchia order bullomorpha cylichna occulta (mighels 1841) class bivalvia subclass pteriomorpha order arcoida bathyarca glacialis (gray 1824) subclass anomalodesmata order pholadomyoida pandora glacialis leach 1819 lyonsia arenosa (möller 1842) total for climatic regions as.. : 6 (16.7%) climatic regions: asb. class gastropoda subclass prosobranchia order neotaenioglossa alvania scrobiculata (möller 1842) alvania jan mayeni (friele 1886) lunatia pallida (broderip & sowerby 1829) order neogastropoda oenopota incisula (verrill 1882) admete viridula (fabricius 1780) subclass opisthobranchia order anaspidea retusa obtusa (montagu 1803) class bivalvia subclass palaeotaxodonta order nuculoida nuculana pernula (müller 1776) yoldia hyperborea lovén 1859 yoldiella lenticula (möller 1842) subclass pteriomorpha order mytiloida musculus laevigatus (gray 1824) musculus niger (gray 1824) crenella decussata (montagu 1803) order pterioida palliolum greenlandicum (sowerby 1842) subclass heterodonta order veneroida axinopsida orbiculata (g.o. sars 1878) tridonta borealis schumacher 1817 tridonta elliptica (brown 1827) clinocardium ciliatum (fabricius 1780) serripes groenlandicus (bruguière 1798) macoma calcarea (gmelin 1791) total for climatic regions asb. : 19 (52.8%) geus bulletin no 3.pmd 28-06-2004, 08:46127 128 climatic regions: asbl class gastropoda subclass prosobranchia order neotaenioglossa natica affinis (gmelin 1790) subclass opisthobranchia order bullomorpha cylichna alba (brown 1827) class bivalvia subclass palaeotaxodonta order nuculoida nuculoma tenuis (montagu 1808) yoldiella frigida (torell 1859) subclass heterodonta order veneroida tridonta montagui (dillwyn 1817) order myoida mya truncata linnaeus 1758 hiatella arctica (linnaeus 1758) total for climatic regions asbl: 7 (19.4%) climatic regions: .sb. class bivalvia subclass pteriomorpha order pterioida chlamys islandica (o.f. müller 1776) subclass heterodonta order myoida panomya arctica (lamarck 1818) total for climatic regions .sb. : 2 (5.6%) climatic regions: .sbl class gastropoda subclass prosobranchia order neotaenioglossa lacuna vincta (montagu 1803) total for climatic regions .sbl: 1 (2.8%) the early/middle weichselian vendsyssel: 36 (14.6%) the mollusc fauna from the vendsyssel region during the early/middle weichselian amounts to 36 species, which would have been even more if not reduced to this number by excluding species regarded as redeposited by v. nordmann (jessen et al. 1910). almost all the recorded species have been found in the skærumhede borings i and ii (jessen et al. 1910; bahnson et al. 1974 respectively), except lunatia pallida, bathyarca glacialis, musculus laevigatus and tridonta borealis, which have been recorded from the older yoldia clay elsewhere in vendsyssel. taken together, the two borings form a most excellent base for evaluating the mollusc faunal development in the early/middle weichselian represented by the arctic sea deposits characterised by portlandia arctica. however, on the basis of the material from skærumhede (bahnson et al. 1974), the portlandia arctica zone is divided into three parts each characterised by other macrofossils from the older to the younger beds: the turritella erosa, balanus crenata and macoma calcarea zones. by doing so, it is emphasised that the development in the arctic part of the marine sequence goes from a deeper-water facies into a shallow-water facies, where in the latter the macoma calcarea species is the dominant bivalve, as it is in the present-day arctic macoma community of east greenland (thorson 1933). from a climatic point of view, nearly all the molluscan species can be found in the high arctic, except lacuna vincta, panomya arctica and chlamys islandica, which are recorded only from the subarctic. however, two of them occur in the turritella erosa zone, which at the same time has the most abundant representation of the other species with northern/arctic affinities, including the purely arctic species portlandia arctica. therefore, no rise in temperature can be suggested on the basis of the molluscan record, only changes in facies through time. together with the dominant molluscan species portlandia arctica and macoma calcarea in the upper part of the arctic sequence – the balanus crenata and the macoma calcarea zones (as seen on fig. 7 in bahnson et al. 1974) – the following mollusc species have been recorded only from these zones: natica affinis, musculus niger, palliolum greenlandicum, and axinopsida orbiculata. all of them can be found in shallow water in the arctic. considering the change of facies from deeper water to shallow water and the lack of climatic changes as seen in the molluscan fauna the question arises of what length of time this development covers. thisquestionhasbeenansweredby the twoamsdates of the topmost part of the marine skærumhede sequence 33 m b.s., which give the age of around 32 000 14c years before present (aar-1410: 32 400 ± 520 and aar-1411: 32 050 ± 420 – both reservoir corrected 14c age (b.p.)). this shows that the marine arctic deposits in this part of the danish area represent nearly the whole part of the early and middle weichselian, because there is an unbroken marine sequence below the level for these ams dates and back into the eemian. on the basis of the correlation of the holmstrup sequence of clay with macoma calcarea, there are reageus bulletin no 3.pmd 28-06-2004, 08:46128 129 sons to think that also the southern part of the kattegat region was part of the older yoldia clay sea far into the weichselian, even though two ams datings from the holmstrup site were infinite (aar-1408: > 38 000 14c age (b.p.) and aar-1409: > 42 000 14c age (b.p.)). skagen age: early/middle weichselian climatic regions: a... class bivalvia subclass palaeotaxodonta order nuculoida portlandia arctica (gray 1824) total for climatic regions a... : 1 (25.0%) climatic regions: asb. class bivalvia subclass palaeotaxodonta order nuculoida nuculana pernula (müller 1776) yoldia hyperborea lovén 1859 subclass pteriomorpha order pterioida palliolum greenlandicum (sowerby 1842) total for climatic regions asb. : 3 (75.0%) the early/middle weichselian skagen: 4 (1.6%) in the skagen region, where all the information coming from the well has been described in more detail earlier, the macrofossil fauna can be presented on the basis of quantitative analyses and sedimentological data (appendix 3). therefore, the material can be seen on the background of a certain bottom community. such a relation was already seen realised in the discussion of the eemian strata in the skagen region, which pointed out that these strata could be correlated with an environment of the deeper part of the present skagerrak. however, the change found in the arctic section of the skagen well on the basis of the mollusc record occurs rather abruptly, turning the scenario into an arctic environment with ice-rafted minerogene material. such a palaeoenvironment is far from the communities demonstrated in recent danish waters, but is well known from east greenland. regarding the finds from the arctic part of the skagen sequence following the temperated eemian strata, it appears that there are no finds of bias from nearshore or indications of other climatic conditions than from the arctic. the four molluscan species found are recorded either only in the high-arctic, viz. portlandia arctica, or in the three zones from the arctic to the boreal, viz. nuculana pernula, yoldiahyperborea, and palliolum greenlandicum. this is a different situation than found in the vendsyssel region, where the two skærumhede wells revealed a clear transition zone – the so-called abra nitida zone – and within the arctic part, with occurrence of molluscs which have been regarded as redeposited (nordmann in jessen et al. 1910), such as: mytilus edulis, pseudamussium septemradiatum, zirfaea crispata, and bittium reticulatum. it has been argued in the present paper in connection with the holocene strata from skagen that, due to the expired isostatic uplift since the weichselian glaciation, the actual depth below present sea level of the holocene beds can be regarded as representing the palaeodepth – taken into account the eustatic movements through time. if this is true, the older strata – the eemian and early weichselian – may also be in a position below present sea level, which could reflect their palaeodepth, here also taking into consideration the older eustatic situation and the question of consolidation. neotectonic movements might be the black horse together with the higher level of the sea during the eemian, max. 7–8 m above the present level. this shows that the transition zone – the abra nitida zone – in the skærumhede ii well is found at a present depth below sea level of between 78 and 88 m (bahnson et al. 1974, fig. 7), while the sharp boundary between the temperate and the arctic zone in the skagen well is at a depth of 180 m b.s. the sea level must have been the same for the two stations at this time, so the difference in depth must be around 100 m. therefore the difference in development of the mollusc fauna within the two sequences depends on depths. this might also explain the occurrences of redeposited molluscs from shallow water in the skærumhede sequence and not in the skagen sequence by a closer coastal environment. the further development of the marine arctic in the skagen region has been truncated by the glaciation within the area. however, as demonstrated earlier, the overlying glacigene deposits have indeed accumulated the ‘missing’ younger marine strata up to an age of around 32 000 before present, as shown by dating of the marine gases. geus bulletin no 3.pmd 28-06-2004, 08:46129 130 late weichselian species sorted after climatic affinities vendsyssel age: late weichselian climatic regions: a... class bivalvia subclass palaeotaxodonta order nuculoida portlandia arctica (gray 1824) subclass heterodonta order veneroida macoma torelli jensen 1904 total for climatic regions a... : 2 (5.7%) climatic regions: as.. class gastropoda subclass opisthobranchia order bullomorpha cylichna occulta (mighels 1841) class bivalvia subclass heterodonta order veneroida macoma loveni jensen 1904 subclass anomalodesmata order pholadomyoida pandora glacialis leach 1819 lyonsia arenosa (möller 1842) total for climatic regions as.. : 4 (11.4%) climatic regions: asb. class gastropoda subclass prosobranchia order neotaenioglossa lunatia pallida (broderip & sowerby 1829) order neogastropoda boreotrophon clathratus (linnaeus 1767) buccinum cyaneum bruguière 1792 neptunea despecta (linnaeus 1758) oenopota turricola (montagu 1803) subclass opisthobranchia order anaspidea retusa obtusa (montagu 1803) class bivalvia subclass palaeotaxodonta order nuculoida nuculana pernula (müller 1776) yoldiella lenticula (möller 1842) subclass pteriomorpha order mytiloida musculus laevigatus (gray 1824) musculus niger (gray 1824) subclass heterodonta order veneroida axinopsida orbiculata (g.o. sars 1878) tridonta borealis schumacher 1817 macoma calcarea (gmelin 1791) total for climatic regions asb. : 13 (37.1%) climatic regions: asbl class gastropoda subclass prosobranchia order neotaenioglossa littorina saxatilis (olivi 1792) natica affinis (gmelin 1790) subclass opisthobranchia order bullomorpha cylichna alba (brown 1827) order thecosomata limacina retroversa (fleming 1823) class bivalvia subclass palaeotaxodonta order nuculoida nuculoma tenuis (montagu 1808) subclass heterodonta order veneroida thyasira flexuosa (montagu 1803) order myoida mya truncata linnaeus 1758 hiatella arctica (linnaeus 1758) total for climatic regions asbl: 8 (22.9%) climatic regions: .sb. class bivalvia subclass pteriomorpha order pterioida chlamys islandica (o.f. müller 1776) subclass heterodonta order veneroida arctica islandica (linnaeus 1767) order myoida zirfaea crispata (linnaeus 1758) total for climatic regions .sb. : 3 (8.6%) climatic regions: .sbl class polyplacophora order neoloricata geus bulletin no 3.pmd 28-06-2004, 08:46130 131 tonicella marmorea (fabricius 1780) class gastropoda subclass prosobranchia order neotaenioglossa lacuna vincta (montagu 1803) order neogastropoda buccinum undatum linnaeus 1758 class bivalvia subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 subclass heterodonta order veneroida macoma balthica (linnaeus 1758) total for climatic regions .sbl: 5 (14.3%) the late weichselian vendsyssel: 35 (14.2%) marine molluscs from the time after the main glaciation of denmark have been recorded only from vendsyssel and skagen, but new studies are increasing our knowledge from the kattegat region, but are not included in this work. the marine mollusc assemblages in vendsyssel have been mainly based on open profiles. the recorded mollusc speciesamount to35.however, therearesomeproblems in classifying the whole fauna in communities. three climatic groups can be demonstrated: a purely arctic and subarctic with six species, an arctic, subarctic and boreal consisting of 13 species, and a group of species which does not enter the arctic but might be found extending into the lusitanian. a fourth group with a wide range can be differentiated according to depth ranges. littorina saxatilis is intertidal, while natica affinis, cylichna alba, and thyasira flexuosa are species found from infratidal to great depths. nuculoma tenuis is found offshore to 300 m, and limacina retroversa is pelagic. so the species with a climatically wide range show the existence of two facies: a littoral and a deeperwater facies. considering the five species with a representation within the subarctic, boreal and lusitanian regions, macoma balthica is a shallow-water species, mytilus edulis is eulittoral and here the species occurs in great quantities, lacuna vincta is intertidal to depths of 60 m, tonicella marmorea from 0 to 183 m but more common at depths of less than 20 m, and finally buccinum undatum which is found sublittorally to great depths (1200 m). the characteristic depth for the dominating part of these molluscs is seen to be the shallow water. two of the species from the subarctic–boreal group, arctica islandica and zirfaea crispata, have very different characteristics as to depth of living. the depth range for arctica islandica is in general intertidal to 480 m, but jensen (1902, pp. 38–39) writes that arctica is a genuine boreal species and bases this on the fact that in the white sea area, which is in the northernmost part of its distribution, it is found in more shallow water than elsewhere. a relatively high temperature is reached only in the shallow water in this region. this means that arctica islandica in the present setting among other molluscs of purely arctic and subarctic relations must be an indicator of boreal waters in the shallow-water environment. zirfaea crispata is on the other hand a clear indicator of shallow water, having its range of depth between the low tide line and out to a depth of about 7 m. the third member of the subarctic–boreal climatic region, chlamys islandica, is known from the tidal zone and down to depths of 300 m. of the 13 species represented within the climatic regions of the arctic, subarctic and boreal, half of the members are infralittoral from 6–10 m to great depths: lunatia pallida, boreotrophon clathratus, neptunea despecta, oenopota turricola, nuculana pernula, yoldiella lenticula, and musculus niger. the other half can be found in the tidal zone but also at greater depths. among these, macoma calcarea and tridonta borealis are the characteristic bivalves in the arctic shallowwater macoma calcarea community with the astarte borealis zone in the most shallow parts from 3 to about 12–14 m in east greenland (thorson 1933, pp. 8–18). according to thorson (1933), the astarte borealis zone is no tide-water community such as for instance the macoma balthica community in some boreal seas. this is discussed in further detail by madsen (1936, p. 71), who concludes: “the littoral fauna [north of c. 66–67°n lat. east greenland] is especially characterised by the absence of littoral molluscs and balanidea, notably mytilus edulis, littorina saxatilis var. groenlandica, and balanus balanoides, all of which occur south of the above-mentioned limit”. the arctic–subarctic species in vendsyssel, represented by four species, can be found at water depths from 2 to 5 m and out to around 200 m, except cylichna occulta, which has a depth range of from 20 m to nearly 400 m. the purely arctic species portlandia arctica and macoma torelli are recorded from 2 and 5 m out to geus bulletin no 3.pmd 28-06-2004, 08:46131 132 around 340 m and 90 m respectively. this shows that the arctic–subarctic part of the recorded species from the vendsyssel region have a wide range of occurrence restricted not by a single species to the shallowwater environment as is the case among the more temperate mollusca and the species with a wide climatic range: arctica islandica, zirfaea crispata, macoma balthica, mytilus edulis, and littorina saxatilis. it can therefore be stated that all the arctic species could be together in deeper water and some also in the more shallow water. however, some of the more temperate species are restricted to shallow water, and arctica islandica is a distinctly shallow-water species in the northern part of its range. so, within the time span of deposition for the late weichselian younger yoldia sea deposits, an amelioration of the shallowwater environment including the tidal zone must have happened. nordmann (1910) noticed that certain marine strata in the vendsyssel region had a distinct littoral fauna, and he suggested that these beds were of a more recent origin than the younger yoldia clay, introducing the so-called zirphaea transgression named after one of the characteristic bivalves from the shallow-water environment discussed above. however, petersen (1984), on the basis of molluscan studies combined with the many 14c dates, could conclude that the deposition of yoldia clay containing a cold marine fauna from deeper water is seen to continue into bølling, but with a contemporaneous temperate fauna in the shallow-water deposits – the zirphaea beds. the evaluation of the marine history from vendsyssel is highly influenced by experiences obtained from work on holocene raised marine deposits and recent bottom samples in east greenland, as seen in petersen (1986b, figs 2, 3), where beds from the astarte borealis zone with mytilus edulis are overlying the ophiocten zone (deeper part of the arctic macoma calcarea community (thorson 1933, pp. 18–27)) with portlandia arctica. these observations may also explain the occurrences of redeposited material mentioned within the discussion of the older yoldia clay at skærumhede in the vendsyssel region. still, it must be regarded as redeposited, but not necessarily differing so much in time, as long as the above-mentioned observations show that a more temperate zone can be found in the shallow-water environment contemporaneous with an arctic fauna in deeper water. skagen age: late weichselian climatic regions: a... class bivalvia subclass palaeotaxodonta order nuculoida portlandia arctica (gray 1824) total for climatic regions a... : 1 (11.1%) climatic regions: as.. class bivalvia subclass pteriomorpha order arcoida bathyarca glacialis (gray 1824) total for climatic regions as.. : 1 (11.1%) climatic regions: asb. class bivalvia subclass palaeotaxodonta order nuculoida nuculana pernula (müller 1776) yoldia hyperborea lovén 1859 yoldiella lenticula (möller 1842) total for climatic regions asb. : 3 (33.3%) climatic regions: asbl class gastropoda subclass opisthobranchia order gymnosomata clione limacina (phipps 1774) class scaphopoda siphonodentalium lobatum (sowerby 1860) class bivalvia subclass palaeotaxodonta order nuculoida yoldiella frigida (torell 1859) total for climatic regions asbl: 3 (33.3%) climatic regions: .sb. class bivalvia subclass palaeotaxodonta order nuculoida nuculana minuta (müller 1776) total for climatic regions .sb. : 1 (11.1%) the late weichselian skagen: 9 (3.6%) the skagen region has contributed with only nine species, out of which the clione limacina species must be taken with some reservation, being based on an geus bulletin no 3.pmd 28-06-2004, 08:46132 133 imprint only. all the recorded species have been taken from the skagen well core superjacent to the older eemian and weichselian deposits discussed above. therefore the fauna represents an assemblage from a certain depth – through time – and can be seen in relation to the sedimentological information (appendix 3). the granulometric composition in the older yoldia sea sequence reflects two maxima on the frequency curve,which tells that part of the material, other than the extremely fine-grained, can be taken as icerafted material. however, no such redeposited material was found in the fauna. the whole mollusc assemblage resembles the arca-astarte crenata community as described by thorson (1934) from hurry inlet, east greenland. first of all the arca glacialis (bathyarca glacialis) is represented among the molluscs recorded from the skagen well. this species is one of the characteristic species from this community. furthermore, the following species are mentioned (thorson 1934, p. 48): siphonodentalium vitreum (s. lobatum), leda pernula (nuculana p.), portlandia arctica, portlandia lenticula (yoldiella l.), portlandia frigida (yoldiella f.), and saxicava arctica (hiatella a.) which have all been recorded from the skagen well. only nuculana minuta, a subarctic–boreal species, and yoldia hyperborea are not recorded from hurry inlet. ockelmann (1958, pp. 19–22) mentioned that nuculana minuta is “lacking in the most high-arctic seas” and that according to thorson (1934) yoldia hyperborea is associated with calm, sheltered places, and besides yoldia hyperborea is otherwise known from only a few places in east greenland. these records from greenland demonstrate a good agreement with the danish late weichselian finds from the skagen well and permit further comparison with the arca-astarte crenata community. according to thorson (1933, p. 67), this community inhabits depths from ca. 45 to ca. 200 m. furthermore, it is poor in species, and the temperature is negative and constant all the year round. the well-dated strata of late weichselian age, from 15 000 – 10 000 before present in 14c years, around 17000 – 11 000 in calibrated age b.p. (see appendix 4) in the skagen well have no influence from the shallow-water boreo-arctic assemblage as found in the vendsyssel region around 13 000 before present (petersen 1984, p. 65), but reveal with their arctic deeperwater mollusc fauna a sharp boundary to the holocene boreo-lusitanian faunas. holocene species sorted after climatic affinities the boundary between the pleistocene and the holocene marine strata as seen in the skagen well is unique within the danish realm. furthermore, the approximately 115 m of holocene marine beds as described earlier represent a well-dated sequence to be compared to the other holocene marine finds recorded in the six regions, many of which offer dated strata as well. the following description will proceed in the same way as taken for the marine pleistocene strata, and it will present the entire molluscan fauna from each region from the point of view of climatic affinities for each species, with subsequent comments upon certain aspects for selected species. within each region, reference to well-dated strata and their molluscan assemblages will be given and worked out to facilitate the correlation in time to assemblages in other regions. finally, the juncture of occurrence during the holocene of some of the molluscan species in all regions can be estimated, as presented in appendix 6. the bælt sea climatic regions: asb. class gastropoda subclass opisthobranchia order anaspidea retusa obtusa (montagu 1803) class bivalvia subclass heterodonta order veneroida tridonta borealis schumacher 1817 total for climatic regions asb. : 2 (4.3%) climatic regions: asbl class gastropoda subclass prosobranchia order neotaenioglossa littorina saxatilis (olivi 1792) lacuna pallidula (da costa 1778) class bivalvia subclass pteriomorpha order mytiloida musculus discors (linnaeus 1767) subclass heterodonta order myoida mya truncata linnaeus 1758 hiatella arctica (linnaeus 1758) geus bulletin no 3.pmd 28-06-2004, 08:46133 134 total for climatic regions asbl: 5 (10.6%) climatic regions: .sb. class bivalvia subclass heterodonta order veneroida arctica islandica (linnaeus 1767) order myoida zirfaea crispata (linnaeus 1758) total for climatic regions .sb. : 2 (4.3%) climatic regions: .sbl class gastropoda subclass prosobranchia order neotaenioglossa littorina obtusata (linnaeus 1758) lacuna vincta (montagu 1803) onoba semicostata (montagu 1803) subclass heterobranchia order heterostropha omalogyra atomus (phillippi 1841) class bivalvia subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 modiolus modiolus (linnaeus 1758) subclass heterodonta order veneroida mysella bidentata (montagu 1803) parvicardium ovale (sowerby 1840) macoma balthica (linnaeus 1758) total for climatic regions .sbl: 9 (19.1%) climatic regions: ..b. class bivalvia subclass heterodonta order myoida mya arenaria linnaeus 1758 total for climatic regions ..b. : 1 (2.1%) climatic regions: ..bl class gastropoda subclass prosobranchia order neotaenioglossa littorina littorea (linnaeus 1758) littorina tenebrosa (montagu 1803) hydrobia ulvae (pennant 1777) hydrobia ventrosa (montagu 1803) rissoa albella lovén 1846 rissoa inconspicua alder 1844 rissoa membranacea (j. adams 1800) bittium reticulatum (da costa 1778) order heterogastropoda triphora adversa (montagu 1803) order neogastropoda hinia reticulata (linnaeus 1758) subclass heterobranchia order heterostropha odostomia conoidea winckworth 1932 subclass opisthobranchia order anaspidea retusa truncatula (bruguière 1792) akera bullata müller 1776 subclass pulmonata order basommatophora lymnaea peregra (müller 1774) class bivalvia subclass pteriomorpha order mytiloida modiolula phaseolina (philippi 1844) order pterioida subclass heterodonta order veneroida parvicardium exiguum (gmelin 1791) parvicardium scabrum (philippi 1844) cerastoderma edule (linnaeus 1758) spisula subtruncata (da costa 1778) angulus tenuis (da costa 1778) scrobicularia plana (da costa 1778) abra alba (wood 1802) paphia aurea (gmelin 1791) tapes decussatus (linnaeus 1758) venerupis pullastra (montagu 1803) order myoida corbula gibba (olivi 1792) barnea candida (linnaeus 1758) total for climatic regions ..bl: 28 (59.6%) the holocene bælt sea: 47 (19.0%) the holocene molluscs in the bælt sea region amount to 47 mainly boreo-lusitanian species. there are no purely lusitanian species, only two arctic–subarctic and boreal species, and five species with a wide range of distribution. this is a clear difference compared to the eemian molluscan fauna from the same region with no less than seven purely lusitanian species. this is not connected with differences in depth ranges. all of the subfossil species from the holocene could be found within the intertidal zone except: odostomia conoidea and parvicardium ovale occurgeus bulletin no 3.pmd 28-06-2004, 08:46134 135 ring below 5–10 m, or infratidal species such as tridonta borealis, modiolus modiolus and abra alba. abra alba characterises the bælt sea deeper-water community and tridonta borealis the community in the baltic, the so-called abra alba and astarte communities sensu petersen (1913, p. 16). the overall dominating part of the species could be associated with the macoma community sensu petersen (1913, p. 14) which nowadays is recorded from the fjords and the more sheltered coasts from the shore and out to a depth of 10–12 m. this community is named after macoma balthica, also present in the bælt sea subfossil molluscan fauna. the great similarity between the subfossil holocene fauna and the recent, as seen from the above-mentioned dates on the communities met with, is also revealed in the present distribution of the species in question. almost all the species occur in the present-day bælt sea or even further into the baltic region – east of darss. the few species no longer found in the bælt sea region are: littorina saxatilis, odostomia conoidea, modiolula phaseolina, ostrea edulis, parvicardium scabrum, angulus tenuis and venerupis pullastra which only extend into the kattegat region today, while paphia aurea, tapes decussatus, and omalogyra atomus are no longer recorded in the danish mollusc fauna. the absence of certain species is most likely an effect of the salinity reached in the bælt sea at present. however, as mentioned by rasmussen (1973, p. 303) in the case of venerupis pullastra, the distribution might be connected with the bottom conditions (see the chapter on molluscan species). most of the species now absent from the bælt sea were present during the atlantic, and, in all, 31 molluscan species have their first appearance in this period (appendix 6). among the 31 species with dated appearances in the atlantic, littorina saxatilis and mya truncata have a wide climatic range. arctica islandica is found in the boreal region, and lacuna vincta, onoba semicostata, mytilus edulis, mysella bidentata, and macoma balthica are subarctic–boreal–lusitanian. the rest of the species appearing in the atlantic amount to 23 species with dated appearances during the atlantic within the bælt sea area. the 31 species constitute 12.6% of the known subfossil molluscan finds from the late quaternary, while the total finds of molluscs from the holocene bælt sea represent 19.0% (47 species out of the 247 subfossil species). this might indicate that the more prolific fauna in the bælt sea was connected with a time interval when the inner danish waters were still affected by a higher tidal impact which expired at the beginning of the subboreal (petersen 1993). however, paphia aurea and ostrea edulis are still met with as food elements, although rather rare, in the iron age ‘køkkenmøddinger’ (kitchen middens) in the western part of the bælt sea (petersen 1985c, fig. 5, p. 22). mya arenaria has been found in the bælt sea embedded in strata from the atlantic. the appearance of this species within the atlantic strata is explained by the deep-burrowing habit of this infauna species. however, the first appearance of the species must be referred to the subatlantic after new dates in the bælt sea region following the study of the immigration of mya arenaria to danish waters (petersen et al. 1992b). the baltic climatic regions: .sbl class bivalvia subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 subclass heterodonta order veneroida macoma balthica (linnaeus 1758) total for climatic regions .sbl: 2 (10.5%) climatic regions: ..bl class gastropoda subclass prosobranchia order neotaenioglossa littorina littorea (linnaeus 1758) littorina tenebrosa (montagu 1803) hydrobia ulvae (pennant 1777) hydrobia ventrosa (montagu 1803) rissoa albella lovén 1846 rissoa inconspicua alder 1844 rissoa membranacea (j. adams 1800) bittium reticulatum (da costa 1778) aporrhais pespelicani (linnaeus 1758) order neogastropoda hinia reticulata (linnaeus 1758) subclass opisthobranchia order anaspidea retusa truncatula (bruguière 1792) subclass pulmonata order basommatophora lymnaea peregra (müller 1774) class bivalvia geus bulletin no 3.pmd 28-06-2004, 08:46135 136 subclass heterodonta order veneroida parvicardium exiguum (gmelin 1791) cerastoderma edule (linnaeus 1758) cerastoderma glaucum (poiret 1798) scrobicularia plana (da costa 1778) order myoida corbula gibba (olivi 1792) total for climatic regions ..bl: 17 (89.5%) the holocene baltic: 19 (7.7%) among the 19 species recorded from the holocene of the baltic, there is relatively many, which do not occur in the recent baltic fauna, viz.: rissoa albella, rissoa inconspicua, bittium reticulatum, hinia reticulata, retusa truncatula, and scrobicularia plana, but they are present in the neighbouring recent bælt sea fauna. furthermore, aporrhais pespelicani can be added which today extends into the kattegat region, while only empty shells have been recorded from the bælt sea region off kiel (arntz et al. 1976). so, although the spectacular oyster and tapes species did not occur as in the bælt sea, the subfossil baltic fauna has quite a few species that no longer live in the baltic. the geological mapping from this region (milthers 1908) does not provide further information on the chronostratigraphic position of these mollusc species, and is also rather poor, since they are based on samples from near present-day sea level. this is also corroborated by munthe (1894, p. 9): “in the south part of the baltic region we possess but comparatively little knowledge of the fauna which results partly from the circumstances that the litorina strata are here to be sought only to a small extent above the sea level”. this has obviously hampered the study within the danish region, so that holocene marine deposits in the baltic are little known up to the present. however, recent activities by marine geologists have given new material from the westernmost part of the baltic – fakse bugt (jensen 1995). as far as the mollusc studies have been submitted (petersen 1994b) but not published in detail, the following comments will be given to fig. 101. the 4 m of sampling from vibrocore 225b comes from the cored section 17.5–13.5 m b.s.l. and has been dated within a time span of 4000 14c years covering the early part of the holocene from the preboreal to the atlantic, through time represented by freshwater, brackish and marine deposits as deduced from the occurrences of mollusc. the loss on ignition shows that the lower third of the sequence has around 25 weight per cent of organic material, while the upper two-thirds of the sequence has less than 5 weight per cent loss on ignition (fig. 101). considering the mollusc species, it appears that the high amount of organic material is not connected solely with the freshwater deposits, but continues into the brackish-water layers. the first rise of sea level is demonstrated by the occurrences of cerastoderma and mytilus and the disappearance of the freshwater molluscs such as valvata macrostoma and the sphaeriidae. the persistent occurrence of the bithynia tentaculata operculae in the oldest part of the brackish-water deposits shows that salinity was lower than 12‰. within this interval, seven mollusc species occur (see fig. 101), viz.: two hydrobia and one rissoa species and the bivalves mytilus, macoma and cerastoderma. the cerastoderma species are rather difficult to identify in all the samples because of their poor state of preservation. at a level of about 15.5 m b.s.l., the change to higher diversity of marine molluscs occurs with such new species as aporrhais pespelicani, nassarius reticulatus and scrobicularia plana besides the steady occurrence of littorina littorea, retusa truncatula and corbula gibba. even bittium reticulatum has been found in one sample. this species tolerates only water with a salinity above 25‰ (sorgenfrei 1958). in all circumstances the more prolific marine fauna is well demonstrated in this part of the sequence, and furthermore, single occurrences of species such as rissoa membranacea, lacuna vincta, and parvicardium exiguum sustain this view. the occurrence of freshwater gastropods is a product of transport from the nearby land, where freshwater streams run into the bay or may be eroded from older deposits by currents. in the topmost part of the core especially fishbones occur up to the present sea bottom at 13.5 m b.s.l. the faunal development in the 4 m core reflects as the oldest element a small lake dated to the time span 9370 ± 135 (k-5649) up to 7900 ± 115 (k-5652) in 14c years b.p. around 7900 b.p. the marine influence is found at the present level of 16.5 m b.s.l. with the establishment of a brackish-water fauna. at a level of 15.30 m b.s.l., which has been dated to 6520 ± 135 14c years b.p. (aar-633), the change to the more prolific littorina sea fauna has taken place. but, the samples geus bulletin no 3.pmd 28-06-2004, 08:46136 137 from higher up in the core do not show any development into the present-day molluscan fauna, first of all because of the absence of mya arenaria that characterises the present-day baltic sea. in conclusion, the littorina sea fauna from the fakse bugt as demonstrated from this core is from the atlantic and shows the molluscs of the transgression around 8000 b.p. (14c years) and the established marine fauna of the atlantic. the fauna from this core contains most of the abovementioned species no longer found in the baltic, and it is shown that this fauna was most probably established during the atlantic. therefore, the marine molluscs from the cored section immigrated to the baltic during the atlantic (appendix 6). among the 15 species with dated appearances in the atlantic only mytilus edulis and macoma balthica have their climatic range within the subarctic–boreal– lusitanian regions, while the rest belong to the boreal– lusitanian group as in the bælt sea area. however, the number of mollusc species from this region is very low with the 19 species found forming only 7.7% of the subfossil fauna. as in the case of the bælt sea, the more prolific fauna can be connected with the atlantic, and this could be caused by the higher tidal amplitude during this time span giving a higher salt content. in the introduction it was mentioned that there is 0 6 12 18 24 30 14 15 16 17 6520 + 135 bp7900 + 115 bp9250 + 130 bp9180 + 130 bp9370 + 135 bpsand gyttja peat lithology li tt or in a lit to re a (l ., 17 58 ) h yd ro bi a ve nt ro sa ( m on ta gu , 1 80 3) h yd ro bi a ul va e (p en na nt , 1 77 7) r is so a al be lla l ov én , 1 84 6 bi tt iu m r et ic ul at um ( da c os ta , 1 77 8) a po rr ha is p es pe lic an i ( l. , 1 75 8) h in ia r et ic ul at a (l ., 17 58 ) r et us a tr un ca tu la ( br ug ui èr e, 1 79 2) m yt ilu s ed ul is l ., 17 58 c er as to de rm a ed ul e (l ., 17 58 ) c er as to de rm a gl au cu m ( po ir et , 1 78 9) m ac om a ba lth ic a (l ., 17 58 ) sc ro bi cu la ri a pl an a (d a c os ta , 1 77 8) c or bu la g ib ba ( o liv i, 17 92 ) 18 –2 0‰ 1 ‰ 10 –2 0‰ 15 –2 0‰ 25 ‰ 18 –2 0‰ 18 –2 0‰ 18 –2 0‰ 6– 10 ‰ 6‰ ~ 6 ‰ 4– 6 ‰ 15 –2 0‰ 15 –2 0‰ ≥ ≥ ≥ ≥ ≥ ≥ ≥ ≥ ≥ ≥ ≥ ≥ ≥ ≥ sp h a er iid a e fo ss il ia v a r ia v al va ta c ri st at a m ül le r, 1 77 4 v al va ta p is ci na lis ( m ül le r, 1 77 4) v al va ta m ac ro st om a st ee nb uc h, 1 84 7 be th yn ia t en ta cu la ta ( l. , 1 75 8) ly m na ea s ta gn al is ( l. , 1 75 8) pl an or bi s co rn eu s (l ., 17 58 ) a ni su s ca ri na tu s (m ül le r, 1 77 4) a ni su s co nt or tu s (l ., 17 58 ) a ni su s cr is ta ( l. , 1 75 8) a ni su s co m pl an at us ( l. , 1 75 8) a ni su s la ev is ( a ld er , 1 83 8) 2‰ 2‰ 0‰ 3– 4 ‰ 5‰ 2‰ 3‰ 5‰ 2– 3 ‰ 4‰ 2– 3 ‰ ≤ ≤ ≤ ≤ ≤ ≤ ≤ ≤ ≤ ≤ ≤ molluscan species with their salinity tolerances loss on ignition weight per cent, %well no, vc225b depth below sea level, m 14c years age the occurrences of molluscan species in the fakse bugt well fig. 101. the occurrences of molluscan species in the fakse bugt well no. vc225b with indication of loss on ignition and sedimentary log with datings (14c years) based on jensen (1995). geus bulletin no 3.pmd 28-06-2004, 08:46137 138 no record of mollusc species from the geological mapping of the island of bornholm (grönwall & milthers 1916). several bottom samples have been analysed from east of bornholm by the author, showing large amounts of tridonta elliptica and tridonta borealis, but the material was never dated, so they could be subrecent specimens. they have therefore not been included in the present list of subfossil holocene molluscs from the baltic. the kattegat climatic regions: asb. class gastropoda subclass opisthobranchia order anaspidea retusa obtusa (montagu 1803) total for climatic regions asb. : 1 (2.2%) climatic regions: asbl class gastropoda subclass prosobranchia order neotaenioglossa littorina saxatilis (olivi 1792) class bivalvia subclass pteriomorpha order mytiloida musculus discors (linnaeus 1767) subclass heterodonta order myoida mya truncata linnaeus 1758 hiatella rugosa (linnaeus 1758) total for climatic regions asbl: 4 (8.9%) climatic regions: .sb. class bivalvia subclass heterodonta order veneroida arctica islandica (linnaeus 1767) total for climatic regions .sb. : 1 (2.2%) climatic regions: .sbl class gastropoda subclass prosobranchia order neotaenioglossa littorina obtusata (linnaeus 1758) lacuna vincta (montagu 1803) skeneopsis planorbis (fabricius 1780) onoba semicostata (montagu 1803) order neogastropoda buccinum undatum linnaeus 1758 subclass heterobranchia order heterostropha omalogyra atomus (phillippi 1841) class bivalvia subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 order pterioida heteranomia squamula (linnaeus 1758) subclass heterodonta order veneroida mysella bidentata (montagu 1803) macoma balthica (linnaeus 1758) total for climatic regions .sbl: 10 (22.2%) climatic regions: ..b. class bivalvia subclass heterodonta order myoida mya arenaria linnaeus 1758 total for climatic regions ..b. : 1 (2.2%) climatic regions: ..bl class gastropoda subclass prosobranchia order archaeogastropoda theodoxus fluviatilis (linnaeus 1758) order neotaenioglossa littorina littorea (linnaeus 1758) littorina tenebrosa (montagu 1803) hydrobia ulvae (pennant 1777) hydrobia ventrosa (montagu 1803) onoba vitrea (montagu 1803) rissoa albella lovén 1846 rissoa inconspicua alder 1844 rissoa membranacea (j. adams 1800) bittium reticulatum (da costa 1778) order heterogastropoda triphora adversa (montagu 1803) order neogastropoda hinia reticulata (linnaeus 1758) subclass heterobranchia order heterostropha brachystomia eulimoides hanley 1844 chrysallida spiralis (montagu 1803) subclass opisthobranchia order anaspidea retusa truncatula (bruguière 1792) akera bullata müller 1776 class bivalvia geus bulletin no 3.pmd 28-06-2004, 08:46138 139 subclass pteriomorpha order pterioida ostrea edulis linnaeus 1758 subclass heterodonta order veneroida acanthocardia echinata (linnaeus 1758) parvicardium exiguum (gmelin 1791) parvicardium scabrum (philippi 1844) cerastoderma edule (linnaeus 1758) cerastoderma glaucum (poiret 1798) scrobicularia plana (da costa 1778) abra alba (wood 1802) paphia aurea (gmelin 1791) tapes decussatus (linnaeus 1758) venerupis pullastra (montagu 1803) order myoida corbula gibba (olivi 1792) total for climatic regions ..bl: 28 (62.2%) the holocene kattegat: 45 (18.2%) the molluscs from the kattegat region have to a great extent been collected in the raised marine forelands to the kattegat itself during the geological mapping in contrast to some of the sampling localities referred to in the preceding two regions. so, although part of the information from djursland comes from borings to a depth of about 10 m b.s.l., it does not present deeperwater deposits as met with in the kattegat proper (petersen 1993). this is also true of the maximum palaeodepth reached, when it is taken into account that the present area is within the isostatic uplift zone with the highest marine shoreline – up to about 10 m – within this region (mertz 1924). forty-five mollusc species have been recorded from the kattegat region. there are no purely lusitanian species, but boreo-lusitanian species. reflecting the above-mentioned facts on maximum palaeodepth, all the species can be found within the tidal/shallow-water zone, viz.: littorina saxatilis, littorina obtusata, littorina tenebrosa, hydrobia ulvae, theodoxus fluviatilis, mytilus edulis, cerastoderma edule, and tapes decussatus. tapes decussatus and paphia aurea are no longer found in danish waters, but occur off western and southern norway in the immediate neighbourhood. rørdam (1891) discussed at some length the different mollusc assemblages in north-eastern sjælland and related these faunas to their relative positions from the open sea and into the innermost part of the fjords. petersen (in rørdam 1891, pp. 106–111), who determined most of the molluscan species in rørdam’s thesis (rørdam 1891, p. 106), stated that: “the deposition of the tapes layers has happened in a period where the danish waters from a hydrographical point of view have been more like the north sea or the open sea than now”. this was clearly demonstrated by the large oyster banks present far into the roskilde fjord in north-eastern sjælland, and it has also been recorded from other mapped areas on fyn and in jylland. the innermost part of the former fjords always carried a rich subfossil mollusc fauna compared to the recent fjord complexes. recent studies from jylland of the faunal changes through time have shown that the rich faunas in the innermost part of such fjord regions developed during the atlantic. this situation is most probably connected with a higher tidal amplitude in the inner danish waters during the atlantic than in the following subboreal period, as argued in petersen (1993). the absence of purely lusitanian species in the midholocene fauna from the kattegat region shows, as mentioned above, that climatic changes have not been of major significance. the most characteristic species for the holocene subfossil fauna in this region – the tapes species – is still to be found within the boreal region. of the 25 dated molluscan species from djursland, 23 species have immigrated during the atlantic along with the transgression (petersen 1993, table 1). only two species, littorina tenebrosa and onoba semicostata, immigrate after the change to more brackish-water conditions that prevailed in the subboreal. among the species with a dated appearance in the atlantic, which form nearly half of the mollusc species known from this region, the dominating climatic group is the boreo-lusitanian, with 18 species. the whole fauna of 45 species from the kattegat region constitutes 18.2% of the subfossil molluscs. this is close to the percentage for the bælt sea area but much less than could be expected regarding the number of species from the recent kattegat region. this recalls the above-mentioned fact that the sampling of the molluscan material has been from raised marine forelands and therefore does not include the deeperwater fauna from the vast area of the kattegat proper. however, as mentioned earlier in connection with the late weichselian marine deposits, new studies are progressing recording the deeper-water fauna with, inter alia, turritella communis from the kattegat during the holocene. geus bulletin no 3.pmd 28-06-2004, 08:46139 140 the limfjord climatic regions: asb. class gastropoda subclass prosobranchia order archaeogastropoda margarites helicinus (phipps 1774) order neogastropoda oenopota turricola (montagu 1803) subclass opisthobranchia order anaspidea retusa obtusa (montagu 1803) class bivalvia subclass heterodonta order veneroida tridonta borealis schumacher 1817 total for climatic regions asb. : 4 (2.7%) climatic regions: asbl class gastropoda subclass prosobranchia order archaeogastropoda acmaea tessulata (müller 1776) order neotaenioglossa littorina saxatilis (olivi 1792) lacuna pallidula (da costa 1778) subclass opisthobranchia order bullomorpha cylichna alba (brown 1827) order anaspidea diaphana minuta brown 1827 class bivalvia subclass palaeotaxodonta order nuculoida nuculoma tenuis (montagu 1808) subclass pteriomorpha order mytiloida musculus discors (linnaeus 1767) subclass heterodonta order veneroida thyasira flexuosa (montagu 1803) order myoida mya truncata linnaeus 1758 hiatella arctica (linnaeus 1758) total for climatic regions asbl: 10 (6.8%) climatic regions: .sb. class bivalvia subclass heterodonta order veneroida arctica islandica (linnaeus 1767) order myoida zirfaea crispata (linnaeus 1758) total for climatic regions .sb. : 2 (1.4%) climatic regions: .sbl class gastropoda subclass prosobranchia order archaeogastropoda acmaea virginea (müller 1776) order neotaenioglossa littorina obtusata (linnaeus 1758) lacuna vincta (montagu 1803) skeneopsis planorbis (fabricius 1780) onoba semicostata (montagu 1803) order neogastropoda nucella lapillus (linnaeus 1758) buccinum undatum linnaeus 1758 subclass heterobranchia order heterostropha omalogyra atomus (phillippi 1841) class bivalvia subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 modiolus modiolus (linnaeus 1758) order pterioida delectopecten vitreus (gmelin 1791) heteranomia squamula (linnaeus 1758) subclass heterodonta order veneroida mysella bidentata (montagu 1803) tellimya ferruginosa (montagu 1803) turtonia minuta (fabricius 1780) parvicardium ovale (sowerby 1840) spisula elliptica (brown 1827) macoma balthica (linnaeus 1758) gari fervensis (gmelin 1791) total for climatic regions .sbl: 19 (12.9%) climatic regions: ..b. class gastropoda subclass heterobranchia order heterostropha chrysallida eximia (jeffreys 1849) total for climatic regions ..b. : 1 (0.7%) climatic regions: ..bl class gastropoda subclass prosobranchia order archaeogastropoda patella vulgata linnaeus 1758 helcion pellucidum (linnaeus 1758) geus bulletin no 3.pmd 28-06-2004, 08:46140 141 iothia fulva (müller 1776) gibbula cineraria (linnaeus 1758) gibbula tumida (montagu 1803) skenea basistriata (jeffreys 1877) order neotaenioglossa littorina littorea (linnaeus 1758) littorina tenebrosa (montagu 1803) lacuna parva (montagu 1803) hydrobia ulvae (pennant 1777) hydrobia ventrosa (montagu 1803) alvania punctura (montagu 1803) cingula semistriata (montagu 1808) onoba vitrea (montagu 1803) rissoa albella lovén 1846 rissoa inconspicua alder 1844 rissoa membranacea (j. adams 1800) rissoa parva (da costa 1779) rissoa violacea desmarest 1814 caecum glabrum (montagu 1803) bittium reticulatum (da costa 1778) turritella communis risso 1826 aporrhais pespelicani (linnaeus 1758) lunatia alderi (forbes 1838) lunatia catena (da costa 1778) order heterogastropoda triphora adversa (montagu 1803) cerithiopsis barleei jeffreys 1867 cerithiopsis tubercularis (montagu 1803) epitonium clathrus (linnaeus 1758) epitonium turtonis (turton 1819) aclis minor (brown 1827) vitreolinaphilippii (rayneval & ponzi 1854) order neogastropoda hinia incrassata (ström 1768) hinia pygmaea (lamarck 1822) hinia reticulata (linnaeus 1758) raphitoma purpurea (montagu 1803) raphitoma linearis (montagu 1803) subclass heterobranchia order heterostropha brachystomia eulimoides hanley 1844 odostomia scalaris macgillivray 1843 chrysallida decussata (montagu 1803) chrysallida indistincta (montagu 1808) chrysallida obtusa (brown 1827) chrysallida spiralis (montagu 1803) ebala nitidissima (montagu 1803) eulimella laevis (brown 1827) eulimella scillae (scacchi 1835) ondina divisa (j. adams 1797) ondina diaphana (jeffreys 1848) odostomia acuta jeffreys 1848 odostomia conoidea winckworth 1932 odostomia turrita hanley 1844 odostomia albella lovén 1846 odostomia plicata (montagu 1803) turbonilla crenata (brown 1827) turbonilla delicata (monterosato 1874) turbonilla lactea (linné 1758) subclass opisthobranchia order bullomorpha acteon tornatilis (linnaeus 1758) cylichna cylindracea (pennant 1777) philine aperta (linnaeus 1767) philine punctata (adams 1800) order anaspidea retusa truncatula (bruguière 1792) retusa umbilicata (montagu 1803) akera bullata müller 1776 class bivalvia subclass palaeotaxodonta order nuculoida nucula nitidosa winckworth 1930 nucula nucleus (linnaeus 1767) subclass pteriomorpha order mytiloida modiolula phaseolina (philippi 1844) modiolus adriaticus (lamarck 1819) modiolaria tumida (hanley 1843) order pterioida aequipecten opercularis (linnaeus 1758) chlamys varia (linnaeus 1758) palliolum striatum (müller 1776) palliolum tigerinum (müller 1776) pododesmus patelliformis (linnaeus 1761) ostrea edulis linnaeus 1758 subclass heterodonta order veneroida lucinoma borealis (linnaeus 1758) lepton nitidum (turton 1822) acanthocardia echinata (linnaeus 1758) parvicardium exiguum (gmelin 1791) parvicardium scabrum (philippi 1844) cerastoderma edule (linnaeus 1758) cerastoderma glaucum (poiret 1798) mactra stultorum (linnaeus 1758) lutraria lutraria (linnaeus 1758) spisula solida (linnaeus 1758) spisula subtruncata (da costa 1778) ensis ensis (linnaeus 1758) phaxas pellucidus (pennant 1777) angulus tenuis (da costa 1778) geus bulletin no 3.pmd 28-06-2004, 08:46141 142 fabulina fabula (gmelin 1791) donax vittatus (da costa 1778) scrobicularia plana (da costa 1778) abra alba (wood 1802) abra nitida (müller 1776) abra prismatica (montagu 1803) chamelea striatula (da costa 1778) clausinella fasciata (da costa 1778) paphia aurea (gmelin 1791) tapes decussatus (linnaeus 1758) timoclea ovata (pennant 1777) venerupis rhomboides (pennant 1777) venerupis pullastra (montagu 1803) mysia undata (pennant 1777) order myoida corbula gibba (olivi 1792) saxicavella jeffreysi winckworth 1930 barnea candida (linnaeus 1758) pholas dactylus linnaeus 1758 subclass anomalodesmata order pholadomyoida thracia phaseolina (lamarck 1818) total for climatic regions ..bl: 107 (72.8%) climatic regions: ...l class gastropoda subclass prosobranchia order archaeogastropoda skenea serpuloides (montagu 1808) order neotaenioglossa alvania lactea (michaud 1830) onoba proxima (forbes & hanley 1850) class bivalvia subclass pteriomorpha order pterioida anomia ephippium linnaeus 1758 total for climatic regions ...l: 4 (2.7%) the holocene limfjord: 147 (59.5%) the mollusc species in the limfjord region amount to 147, and within this relatively high number only four purely lusitanian species are found. skenea serpuloides is known from recent waters in the british isles and southward, alvania lactea and onoba proxima from the western coast of britain and to the south, and anomia ephippium also from the british isles, although including the orkney islands and south to the mediterranean. according to jensen & spärck (1934), anomia ephippium has often not been separated from heteranomia squamula. on the occurrences of anomia ephippium and the three lusitanian gastropods, which are all tiny and difficult species to work with in the subfossil state, the question may then arise whether much reliance should be put in these circumstances indicating that the subfossil holocene fauna of the limfjord region had a more lusitanian affinity than is the case at present. the next climatic group is the boreo-lusitanian, which is by far the largest, with 107 species, 72.8% of all the species recorded from this region. in the limfjord region it is possible to distinguish between the recent fauna that arrived after the breakthrough at the agger tange in 1825, when the limfjord again was established as a saltwater basin, and the older long stage before the middle ages and back to the transgression in the early holocene. the faunal record of the recent mollusc species from the limfjord is an illustration of how fast a population can be established, although not studied from the very beginning (spärck 1943, p. 78). the 85 species established there during the period of about 100 years picture recent immigration (petersen 1986a, p. 223). already the early studies by collin (1884) and petersen (1888) demonstrated that within the subfossil holocene fauna there was a certain number of mollusc species no longer known from the present fauna in the limfjord or within the danish waters at all. these deposits were called the tapes beds by petersen (1888, p. 56). the tapes species sensu petersen (1888) include: paphia aurea, tapes decussatus, and venerupis rhomboides, the last one has lately been recorded as part of the recent danish fauna (jensen & knudsen 1995). the way to have a firmer basis for discriminating between the recent and the holocene fauna in the limfjord region would be to split the boreal climatic region into the three zones: the high-boreal, the midboreal and the low-boreal following the indications on distribution given in the chapter on the molluscan species, following feyling-hanssen (1955) in fig. 4. in this way it turns out that the recent boreo-lusitanian group, with a number of 51, has 45% reaching to the north into the high-boreal, 47% the mid-boreal, and 9% the low-boreal zone. while the subfossil boreolusitanian group of 107 species shows 41% reaching into the high-boreal, 45% the mid-boreal, and 14% the low-boreal zone. there is, so to say, only a slightly higher affinity to more temperate southern waters for the subfossil molluscan fauna, meaning that a hypsothermal period is not clearly demonstrated in the mageus bulletin no 3.pmd 28-06-2004, 08:46142 143 rine environment from the limfjord. in the case of the limfjord region, it should be emphasised that we do have a special situation in connection with the living depths of certain species. taking into consideration again the largest group of molluscs recorded from the holocene limfjord deposits – the boreo-lusitanian group – 12 species have their main occurrence below the tidal zone, and 26 species out of the 107 species encountered occur at a deeper level. when this is seen from the fact that the collection of molluscs during the geological mapping has been done mostly in outcrops, and that the highest marine limit goes up to only about 5 m a.s.l. in this area (mertz 1924), the palaeodepth reached cannot be as deep as figured on the basis of the general depth range of the recent molluscs in danish waters. this was a point stressed by nordmann (jessen 1905, pp. 151–152) and clearly indicates a smaller living depth, found for the molluscs from the limfjord region than from other regions in denmark during the holocene. such a difference in the habitat can still be observed on the beaches of the limfjord, where inter alia the deeper-living mytilid modiolus modiolus is found washed ashore quite commonly in quantities not seen elsewhere along the shores of denmark. the remaining climatic groups are the ones with a distribution extending into the arctic or the subarctic, some with a wide range reaching the lusitanian or the boreal zones to the south. here it should be pointed out that all of the species with their southern limit in the boreal zone have been recorded also from the low-boreal sector, although tridonta borealis is said to be rare in the north sea, but is common in the bælt sea and baltic regions. from this it can be concluded that all the species recorded from the holocene/recent limfjord region could have coexisted in various habitats – except for the three purely lusitanian species. several 14c dates from the limfjord region on molluscs from the holocene reveal the immigration time of a few species (petersen & rasmussen 1995a, table 1), although the actual number of dated shells of certain species are low. from the atlantic: mytilus edulis, cerastoderma edule, ostrea edulis, arctica islandica, spisula subtruncata, corbula gibba, scrobicularia plana, acanthocardia echinata, and venerupis pullastra. from the subboreal: lucinoma borealis, tapes decussatus, and paphia aurea. from the subatlantic: donax vittatus. this is a function of the same species being used in many more datings, because specimens of these species have been present in a sufficient number or weight to allow a conventional 14c dating of the whole assemblage or bed of molluscs. therefore, also the other species present in the samples or stratum shall be considered dated, just as dating of certain levels in borings are taken into account (petersen 1976, 1981, 1985b, 1986c; rasmussen & petersen 1980). in this way a far higher number of first occurrence of species can be demonstrated, still on the basis of absolute dates, as presented in appendix 6. for as many as 114 species out of the total number of recorded species (147) from the holocene mollusc faunas in the limfjord area first occurrence has been dated: 77 species dated to the atlantic, 36 species to the subboreal and one species to the subatlantic. in the atlantic and the subboreal the dominating climatic groups are the boreal–lusitanian species with 55 and 28 species, forming 71.4% and 77.8% respectively. two of the purely lusitanian species which have been dated, alvania lactea and onoba proxima, appeared in the subboreal. the total number of species recorded from the limfjord (147) constitutes 59.5% of the subfossil late quaternary molluscs, which is much higher than seen in the inner danish waters. the north sea climatic regions: asb. class gastropoda subclass opisthobranchia order anaspidea retusa obtusa (montagu 1803) total for climatic regions asb. : 1 (1.1%) climatic regions: asbl class gastropoda subclass prosobranchia order neotaenioglossa littorina saxatilis (olivi 1792) lacuna pallidula (da costa 1778) subclass opisthobranchia order bullomorpha cylichna alba (brown 1827) class bivalvia subclass palaeotaxodonta order nuculoida nuculoma tenuis (montagu 1808) subclass pteriomorpha geus bulletin no 3.pmd 28-06-2004, 08:46143 144 order mytiloida musculus discors (linnaeus 1767) subclass heterodonta order veneroida thyasira flexuosa (montagu 1803) order myoida mya truncata linnaeus 1758 hiatella arctica (linnaeus 1758) total for climatic regions asbl: 8 (8.4%) climatic regions: .sb. class bivalvia subclass heterodonta order veneroida arctica islandica (linnaeus 1767) order myoida zirfaea crispata (linnaeus 1758) total for climatic regions .sb. : 2 (2.1%) climatic regions: .sbl class gastropoda subclass prosobranchia order neotaenioglossa littorina obtusata (linnaeus 1758) lacuna vincta (montagu 1803) order neogastropoda nucella lapillus (linnaeus 1758) buccinum undatum linnaeus 1758 class bivalvia subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 order pterioida heteranomia squamula (linnaeus 1758) subclass heterodonta order veneroida mysella bidentata (montagu 1803) tellimya ferruginosa (montagu 1803) parvicardium ovale (sowerby 1840) spisula elliptica (brown 1827) macoma balthica (linnaeus 1758) total for climatic regions .sbl: 11 (11.6%) climatic regions: ..b. class gastropoda subclass prosobranchia order neotaenioglossa cingula turgida (jeffreys 1870) total for climatic regions ..b. : 1 (1.1%) climatic regions: ..bl class gastropoda subclass prosobranchia order archaeogastropoda gibbula cineraria (linnaeus 1758) theodoxus fluviatilis (linnaeus 1758) order neotaenioglossa littorina littorea (linnaeus 1758) lacuna parva (montagu 1803) hydrobia ulvae (pennant 1777) hydrobia ventrosa (montagu 1803) onoba vitrea (montagu 1803) rissoa albella lovén 1846 rissoa inconspicua alder 1844 rissoa membranacea (j. adams 1800) rissoa violacea desmarest 1814 caecum glabrum (montagu 1803) bittium reticulatum (da costa 1778) turritella communis risso 1826 aporrhais pespelicani (linnaeus 1758) lunatia alderi (forbes 1838) lunatia catena (da costa 1778) order heterogastropoda triphora adversa (montagu 1803) epitonium clathrus (linnaeus 1758) aclis ascaris (turton 1819) aclis minor (brown 1827) aclis walleri jeffreys 1867 order neogastropoda hinia pygmaea (lamarck 1822) hinia reticulata (linnaeus 1758) subclass heterobranchia order heterostropha brachystomia eulimoides hanley 1844 chrysallida indistincta (montagu 1808) chrysallida spiralis (montagu 1803) eulimella laevis (brown 1827) ondina diaphana (jeffreys 1848) odostomia conoidea winckworth 1932 odostomia albella lovén 1846 turbonilla crenata (brown 1827) turbonilla delicata (monterosato 1874) turbonilla lactea (linné 1758) subclass opisthobranchia order bullomorpha acteon tornatilis (linnaeus 1758) order anaspidea retusa truncatula (bruguière 1792) retusa umbilicata (montagu 1803) class bivalvia subclass palaeotaxodonta order nuculoida nucula nitidosa winckworth 1930 geus bulletin no 3.pmd 28-06-2004, 08:46144 145 nucula nucleus (linnaeus 1767) nucula sulcata (bronn 1831) subclass pteriomorpha order pterioida chlamys varia (linnaeus 1758) ostrea edulis linnaeus 1758 subclass heterodonta order veneroida lepton nitidum (turton 1822) acanthocardia echinata (linnaeus 1758) parvicardium exiguum (gmelin 1791) parvicardium scabrum (philippi 1844) parvicardium minimum (philippi 1836) cerastoderma edule (linnaeus 1758) mactra stultorum (linnaeus 1758) spisula solida (linnaeus 1758) spisula subtruncata (da costa 1778) ensis ensis (linnaeus 1758) phaxas pellucidus (pennant 1777) angulus tenuis (da costa 1778) fabulina fabula (gmelin 1791) donax vittatus (da costa 1778) scrobicularia plana (da costa 1778) abra alba (wood 1802) abra nitida (müller 1776) chamelea striatula (da costa 1778) clausinella fasciata (da costa 1778) paphia aurea (gmelin 1791) tapes decussatus (linnaeus 1758) timoclea ovata (pennant 1777) venerupis pullastra (montagu 1803) dosinia exoleta (linnaeus 1758) dosinia lincta (montagu 1803) order myoida corbula gibba (olivi 1792) saxicavella jeffreysi winckworth 1930 barnea candida (linnaeus 1758) subclass anomalodesmata order pholadomyoida cochlodesma praetenue (pulteney 1799) thracia phaseolina (lamarck 1818) total for climatic regions ..bl: 72 (75.8%) the holocene north sea: 95 (38.5%) from the north sea region, 95 species have been recorded, and here no purely lusitanian mollusc species has been found. almost all the north sea finds are recorded also from the limfjord. the boreo-lusitanian group in the north sea region is also, compared to the limfjord, by far the largest group, with 72 species. in this group, only five species can be pointed out as not occurring in the holocene of the limfjord, viz.: parvicardium minimum, clausinella fasciata, dosinia exoleta, dosinia lincta, and cochlodesma praetenue. furthermore, donax vittatus, as remarked earlier in the chapter on the mollusc species, this species does not belong to the limfjord proper, but is recorded from old (subatlantic) beach ridges once facing the skagerrak. today, this bivalve is bound to the exposed coast of denmark, not penetrating into the inner danish waters. the dated occurrences of this species in the north sea area fall in the subatlantic, when the present coastline of jylland was developed (petersen 1994a). donax vittatus has a distribution to the north up into the mid-boreal sector between trondheim fjord and lofoten. the other boreo-lusitanian species in the north sea region which are not in the limfjord go as far north as the high-boreal sector north of lofoten. it is seen that the species not found in the limfjord but in the holocene north sea are to be regarded not as newcomers showing any amelioration, but more probably as representing other conditions prevailing in the north sea area than in the limfjord, as seen in the case of donax vittatus. parvicardium minimum is common only at depths of more than 30 m. clausinella fasciata is far from common in the inner danish waters and, when occurring, is so only at depths of between 15 and 30 m. dosinia exoleta and dosinia lincta can be found from the intertidal zone and out to depths of 70 and 200 m respectively; these species have been recorded from the northern kattegat, although the latter extends into the øresund (jensen & knudsen 1995). the great similarity and the few differences found when the holocene north sea fauna is compared with the limfjord record are explained in the recent study on the agger tange complex (petersen 1994a, 1998). this study, being based on material from several vibrocores west of the agger tange complex, revealed that the jydske rev forms a continuation of the limfjord complex 75 km further towards the west. ams dates of molluscs from this part of the north sea and the oldest cored sections in the agger tange complex show that the marine record can be established from the preboreal and up into the subatlantic (petersen 1985a). the fauna dated from the older part of the holocene is very similar to the limfjord fauna and the younger fauna, containing only a few differences to geus bulletin no 3.pmd 28-06-2004, 08:46145 146 the limfjord fauna, as indicated by the dosinia species. during the older stages of the holocene, the jydske rev complex formed a landscape much the same as that of the present limfjord, while in the younger part of the holocene erosion has taken place and the present coastline of western jylland developed. the appearance of such molluscs as the dosinia and donax species must be connected with the new facies in the exposed coastal areas rather than indicating climatic changes. the well-dated molluscs from the vibrocores in the north sea and borings in the agger tange complex allow the fixing of a first appearance of most of the 95 species recorded from the north sea region as seen in appendix 6 (petersen 1985a, 1994a). the total number of recorded species from the holocene north sea is 95, forming 38.5% of the subfossil molluscs from the late quaternary, which is less than recorded from the limfjord, but out of this number 80 species have been dated with their first appearance. what is of special interest from this area is that the record also covers 26 dated species from the preboreal–boreal and furthermore 27 species from the atlantic, 19 species from the subboreal and eight species from the subatlantic. in all time intervals, the boreo-lusitanian group is by far the dominating part as shown in table 1. as the boreo-lusitanian group in the north sea constitutes the most temperate species, the atlantic can be pointed out as having a slightly higher proportion of warmer mollusc faunal element than the other periods. in order to find the climatic trend for the holocene, the climatic affinity of the recent species not found in the subfossil deposits is compared with that of the subfossil holocene species (table 2). when the affinities to the climatic regions for the 183 holocene subfossil species are compared with those of the 94 recent species that do not occur in the holocene subfossil fauna, a slightly higher affinity to the more temperate regions for the subfossil species appears. from the north sea material it was concluded that the atlantic has a slightly higher proportion of warmer elements than the other periods. so within the danish realm the indication of amelioration, as seen from the molluscan material, points to the atlantic. vendsyssel climatic regions: asb. class gastropoda subclass prosobranchia order neogastropoda oenopota turricola (montagu 1803) subclass opisthobranchia order anaspidea retusa obtusa (montagu 1803) class bivalvia subclass heterodonta order veneroida macoma calcarea (gmelin 1791) total for climatic regions asb. : 3 (2.3%) climatic regions: asbl class gastropoda subclass prosobranchia order neotaenioglossa littorina saxatilis (olivi 1792) lacuna pallidula (da costa 1778) subclass opisthobranchia order anaspidea diaphana minuta brown 1827 class bivalvia subclass pteriomorpha order mytiloida musculus discors (linnaeus 1767) subclass heterodonta order veneroida thyasira flexuosa (montagu 1803) asb. 5 (0.7%) 5 (5.4%) asbl 11 (6.0%) 9 (9.6%) .sb. 3 (1.6%) 0 .sbl 19 (10.4%) 7 (7.4%) ..b. 3 (1.6%) 7 (7.4%) ..bl 136 (74.3%) 64 (68.1%) ...l 6 (3.3%) 2 (2.1%) climatic regions holocene subfossil spp. recent spp. table 2. the climatic trend for the holocene geus bulletin no 3.pmd 09-07-2004, 09:10146 147 order myoida mya truncata linnaeus 1758 hiatella arctica (linnaeus 1758) hiatella rugosa (linnaeus 1758) total for climatic regions asbl: 8 (6.0%) climatic regions: .sb. class bivalvia subclass heterodonta order veneroida arctica islandica (linnaeus 1767) order myoida zirfaea crispata (linnaeus 1758) total for climatic regions .sb. : 2 (1.5%) climatic regions: .sbl class gastropoda subclass prosobranchia order archaeogastropoda acmaea virginea (müller 1776) order neotaenioglossa littorina obtusata (linnaeus 1758) lacuna vincta (montagu 1803) skeneopsis planorbis (fabricius 1780) onoba semicostata (montagu 1803) order neogastropoda nucella lapillus (linnaeus 1758) buccinum undatum linnaeus 1758 subclass heterobranchia order heterostropha omalogyra atomus (phillippi 1841) class bivalvia subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 modiolus modiolus (linnaeus 1758) order pterioida heteranomia squamula (linnaeus 1758) subclass heterodonta order veneroida mysella bidentata (montagu 1803) tellimya ferruginosa (montagu 1803) turtonia minuta (fabricius 1780) parvicardium ovale (sowerby 1840) spisula elliptica (brown 1827) macoma balthica (linnaeus 1758) gari fervensis (gmelin 1791) total for climatic regions .sbl: 18 (13.5%) climatic regions: ..b. class bivalvia subclass heterodonta order myoida mya arenaria linnaeus 1758 total for climatic regions ..b. : 1 (0.8%) climatic regions: ..bl class gastropoda subclass prosobranchia order archaeogastropoda helcion pellucidum (linnaeus 1758) gibbula cineraria (linnaeus 1758) gibbula tumida (montagu 1803) order neotaenioglossa littorina littorea (linnaeus 1758) littorina tenebrosa (montagu 1803) lacuna parva (montagu 1803) hydrobia ulvae (pennant 1777) alvania cimicoides (forbes 1844) alvania punctura (montagu 1803) cingula semistriata (montagu 1808) onoba vitrea (montagu 1803) rissoa albella lovén 1846 rissoa inconspicua alder 1844 rissoa membranacea (j. adams 1800) rissoa parva (da costa 1779) rissoa violacea desmarest 1814 caecum glabrum (montagu 1803) bittium reticulatum (da costa 1778) turritella communis risso 1826 aporrhais pespelicani (linnaeus 1758) lunatia alderi (forbes 1838) lunatia catena (da costa 1778) order heterogastropoda triphora adversa (montagu 1803) epitonium clathrus (linnaeus 1758) epitonium turtonis (turton 1819) vitreolina philippii (rayneval & ponzi 1854) order neogastropoda neptunea antiqua (linnaeus 1758) hinia incrassata (ström 1768) hinia pygmaea (lamarck 1822) hinia reticulata (linnaeus 1758) cytharella coarctata (forbes 1840) raphitoma linearis (montagu 1803) subclass heterobranchia order heterostropha brachystomia eulimoides hanley 1844 odostomia scalaris macgillivray 1843 chrysallida indistincta (montagu 1808) chrysallida obtusa (brown 1827) chrysallida spiralis (montagu 1803) ebala nitidissima (montagu 1803) geus bulletin no 3.pmd 09-07-2004, 09:10147 148 eulimella laevis (brown 1827) odostomia conoidea winckworth 1932 odostomia turrita hanley 1844 odostomia albella lovén 1846 odostomia plicata (montagu 1803) turbonilla lactea (linné 1758) subclass opisthobranchia order bullomorpha acteon tornatilis (linnaeus 1758) cylichna cylindracea (pennant 1777) philine aperta (linnaeus 1767) philine punctata (adams 1800) order anaspidea retusa truncatula (bruguière 1792) retusa umbilicata (montagu 1803) akera bullata müller 1776 class bivalvia subclass palaeotaxodonta order nuculoida nucula nitidosa winckworth 1930 nucula nucleus (linnaeus 1767) subclass pteriomorpha order mytiloida modiolula phaseolina (philippi 1844) modiolus adriaticus (lamarck 1819) modiolaria tumida (hanley 1843) order pterioida aequipecten opercularis (linnaeus 1758) chlamys varia (linnaeus 1758) pecten maximus (linnaeus 1758) pododesmus patelliformis (linnaeus 1761) ostrea edulis linnaeus 1758 subclass heterodonta order veneroida lucinoma borealis (linnaeus 1758) lepton nitidum (turton 1822) kellia suborbicularis (montagu 1803) acanthocardia echinata (linnaeus 1758) parvicardium exiguum (gmelin 1791) parvicardium scabrum (philippi 1844) cerastoderma edule (linnaeus 1758) laevicardium crassum (gmelin 1791) mactra stultorum (linnaeus 1758) lutraria lutraria (linnaeus 1758) spisula solida (linnaeus 1758) spisula subtruncata (da costa 1778) ensis ensis (linnaeus 1758) phaxas pellucidus (pennant 1777) angulus tenuis (da costa 1778) fabulina fabula (gmelin 1791) donax vittatus (da costa 1778) gari depressa (pennant 1777) scrobicularia plana (da costa 1778) abra alba (wood 1802) abra nitida (müller 1776) abra prismatica (montagu 1803) chamelea striatula (da costa 1778) clausinella fasciata (da costa 1778) paphia aurea (gmelin 1791) tapes decussatus (linnaeus 1758) timoclea ovata (pennant 1777) venerupis rhomboides (pennant 1777) venerupis pullastra (montagu 1803) dosinia exoleta (linnaeus 1758) dosinia lincta (montagu 1803) mysia undata (pennant 1777) order myoida corbula gibba (olivi 1792) saxicavella jeffreysi winckworth 1930 barnea candida (linnaeus 1758) pholas dactylus linnaeus 1758 subclass anomalodesmata order pholadomyoida thracia phaseolina (lamarck 1818) total for climatic regions ..bl: 98 (73.7%) climatic regions: ...l class gastropoda subclass prosobranchia order neotaenioglossa alvania lactea (michaud 1830) trivia monacha (da costa 1778) class bivalvia subclass pteriomorpha order pterioida anomia ephippium linnaeus 1758 total for climatic regions ...l: 3 (2.3%) the holocene vendsyssel: 133 (53.8%) in the vendsyssel region the mollusc species amount to 133, nearly the same as recorded from the limfjord region (147). however, some other species occur, although the grouping of species according to the climatic regions to which they belong is almost the same. this is a very important fact considering the large number of species recorded from both regions and according to the conclusions drawn from the limfjord material when compared to the recent fauna in the limfjord, with only a slight difference between the climatic affinities of the subfossil and the recent molgeus bulletin no 3.pmd 09-07-2004, 09:10148 149 luscs, including the recent species which have invaded the limfjord since 1825, but not recorded in the subfossil material. the vendsyssel fauna counts three purely lusitanian elements, but they are far from common, and alvania lactea and trivia monacha are not typical for the former vendsyssel palaeoenvironment. within the boreo-lusitanian group, only ten molluscs are new in the fauna compared to the holocene of the limfjord: alvania cimicoides, neptunea antiqua, cytharella coarctata, pecten maximus, kellia suborbicularis, laevicardium crassum, gari depressa, clausinella fasciata, dosinia exoleta, and dosinia lincta. nordmann (1904) states that the mollusc fauna recorded from the sites north of frederikshavn have a distinct lusitanian affinity. however, as seen above, only ten species are new to the vendsyssel region compared to the limfjord, and here four species reach into the high-boreal sector (north of lofoten), five species into the mid-boreal (between lofoten and trondheim), and only one, neptunea antiqua, has its northern limit within the low-boreal sector (south of trondheim and to the channel). therefore, the designation of the dosinia fauna cannot be one of the special southern appearances. however, nordmann also states that the same species in the dosinia fauna are new compared to the tapes fauna. this is true only for four species, viz. laevicardium crassum, neptunea antiqua, trivia monacha, and kellia suborbicularis, while venerupis rhomboides, lutraria lutraria, gari fervensis and hinia incrassata also occur to the south of frederikshavn and in the limfjord region. furthermore, the characterising species of these beds – dosinia exoleta – is now also recorded from the subatlantic beds at a depth of 32.5 m in the north sea, occurring together with, among others, dosinia lincta. nordmann (1904, pp. 30–31) also argued for the dosinia fauna to be a shallow-water assemblage and discussed in greater detail dosinia exoleta, venerupis rhomboides, lutraria lutraria, laevicardium crassum, lucinoma borealis, arctica islandica, and neptunea antiqua. from the general information at hand on depth relations of these species (chapter on molluscan species), it appears that all except neptunea antiqua can be found from the intertidal zone and out to various depths in deeper water, while neptunea antiqua has a minimum depth of 15 m. taking into account all the recorded molluscan finds from the vendsyssel region, it appears that 84 species (63%) can be found in the tidal–intertidal zone and 48 species (36%) in water deeper than that. therefore it cannot be characteristic of the dosinia fauna that it is a shallow-water assemblage. the new dates which have been used in the discussion of the dosinia fauna date its appearance by the oldest date for the dosinia exoleta, at 4240 b.p. (k5318; in petersen 1991b). skagen climatic regions: asb. class gastropoda subclass prosobranchia order neogastropoda oenopota turricola (montagu 1803) total for climatic regions asb. : 1 (1.4%) climatic regions: asbl class gastropoda subclass prosobranchia order neotaenioglossa lacuna pallidula (da costa 1778) subclass opisthobranchia order bullomorpha cylichna alba (brown 1827) class bivalvia subclass pteriomorpha order mytiloida musculus discors (linnaeus 1767) subclass heterodonta order veneroida thyasira flexuosa (montagu 1803) order myoida hiatella arctica (linnaeus 1758) total for climatic regions asbl: 5 (7.0%) climatic regions: .sb. class bivalvia subclass palaeotaxodonta order nuculoida nuculana minuta (müller 1776) subclass heterodonta order veneroida arctica islandica (linnaeus 1767) total for climatic regions .sb. : 2 (2.8%) climatic regions: .sbl class gastropoda subclass prosobranchia order neogastropoda buccinum undatum linnaeus 1758 geus bulletin no 3.pmd 09-07-2004, 09:10149 150 class bivalvia subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 order pterioida heteranomia squamula (linnaeus 1758) subclass heterodonta order veneroida mysella bidentata (montagu 1803) tellimya ferruginosa (montagu 1803) turtonia minuta (fabricius 1780) gari fervensis (gmelin 1791) total for climatic regions .sbl: 7 (9.9%) climatic regions: ..b. class bivalvia subclass heterodonta order myoida mya arenaria linnaeus 1758 total for climatic regions ..b. : 1 (1.4%) climatic regions: ..bl class gastropoda subclass prosobranchia order neotaenioglossa hydrobia ulvae (pennant 1777) barleeia unifasciata (montagu 1803) onoba vitrea (montagu 1803) rissoa albella lovén 1846 rissoa violacea desmarest 1814 bittium reticulatum (da costa 1778) turritella communis risso 1826 aporrhais pespelicani (linnaeus 1758) lunatia alderi (forbes 1838) lunatia montagui (forbes 1838) order heterogastropoda epitonium trevelyanum (johnston 1841) aclis minor (brown 1827) polygireulima sinuosa (sacco 1836) vitreolina philippii (rayneval & ponzi 1854) graphis albida (kanmacher 1798) melanella lubrica (monterosato 1891) melanella alba (da costa 1778) hemiaclis ventrosa (jeffreys ms fricle 1874) order neogastropoda hinia pygmaea (lamarck 1822) hinia reticulata (linnaeus 1758) mangelia brachystoma (philippi 1844) subclass heterobranchia order heterostropha chrysallida decussata (montagu 1803) eulimella scillae (scacchi 1835) odostomia conoidea winckworth 1932 odostomia umbilicaris (malm 1863) turbonilla delicata (monterosato 1874) turbonilla sinuosa (jeffreys 1884) subclass opisthobranchia order anaspidea retusa truncatula (bruguière 1792) retusa umbilicata (montagu 1803) class bivalvia subclass palaeotaxodonta order nuculoida nucula nitidosa winckworth 1930 nucula nucleus (linnaeus 1767) subclass pteriomorpha order pterioida chlamys varia (linnaeus 1758) ostrea edulis linnaeus 1758 subclass heterodonta order veneroida acanthocardia echinata (linnaeus 1758) parvicardium minimum (philippi 1836) mactra stultorum (linnaeus 1758) spisula subtruncata (da costa 1778) phaxas pellucidus (pennant 1777) angulus tenuis (da costa 1778) fabulina fabula (gmelin 1791) tellina pygmaea (lovén 1846) donax vittatus (da costa 1778) abra alba (wood 1802) abra nitida (müller 1776) abra prismatica (montagu 1803) chamelea striatula (da costa 1778) timoclea ovata (pennant 1777) order myoida corbula gibba (olivi 1792) saxicavella jeffreysi winckworth 1930 barnea candida (linnaeus 1758) pholas dactylus linnaeus 1758 subclass anomalodesmata order pholadomyoida lyonsia norvegica (gmelin 1791) cochlodesma praetenue (pulteney 1799) thracia phaseolina (lamarck 1818) total for climatic regions ..bl: 54 (76.1%) climatic regions: ...l class gastropoda subclass prosobranchia order heterogastropoda vitreolina collensi (sykes 1903) geus bulletin no 3.pmd 09-07-2004, 09:10150 151 total for climatic regions ...l: 1 (1.4%) the holocene skagen: 71 (28.7%) from the skagen boring, 71 holocene species have been recorded. the boreo-lusitanian group dominates with 54 species (76.1%) of the holocene mollusc species from the skagen well. one purely lusitanian species, vitreolina collensi, has been found, while all the other species occur in the boreal and to some extent the arctic. in this way the skagen well material resembles that of other regions like the vendsyssel, limfjord and north sea during the holocene. the skagen well material has all been recorded to certain stratigraphical levels, as seen in appendix 6, so the climatic indications through time appear, but the number of molluscan species is very low. in the preboreal/boreal, only three species have been recorded. this has been explained as a result of a deeper water where the echinoids dominate. higher up in the sequence, the number of molluscs increases – 23 species in the subboreal and 68 species in the subatlantic. through the chronostratigraphical levels, the climatic regions of the boreo-lusitanian from the dominating one, and the purely lusitanian vitreolina collensi as mentioned above occurs in the subboreal and subatlantic. however, as already stated, the development of the facies in the skagen well during the holocene does change the environment from the deeper-water facies with few molluscs through the bottom community with turritella communis into the prolific shallow-water community. in this way the youngest part covering the subboreal–subatlantic is also by far the part with the highest species diversity. the environmental changes within the seven regions through the late quaternary evaluated by the molluscan communities met with in the seven stages the seven chronological stages which have been described according to their climatic affinities are seen in fig. 102 and fig. 103 covering the eemian, the early/ middleweichselian, thelateweichselian, thepreboreal/ boreal, the atlantic, the subboreal, and the subatlantic. in this way the climatic cycle during the late quaternary is demonstrated on the basis of marine mollusc species which indicate that the eemian has by far the highest amount of the more temperate species, while the holocene reached its maximum during the atlantic, although only slightly more than the other stages within the holocene, as already commented upon in the previous chapter. it is generally accepted that the eemian summer temperatures were higher – about 2°c above the present. with glaciers smaller than the present day, this means that the sea level was 4–6 m higher than today (andersen & borns 1994, pp. 44–49). and as pointed out by donner (1995, p. 39): “the submergence was clearly greater after the saalian glaciation than after the weichselian and possibly after the older glaciations”. donner sees this in northern europe as “a result of a comparatively great downwarping of the earth’s crust during the extensive saalian glaciation”. the rebound since the last glaciation has come to an end within the danish area (petersen 1985c, 1991b). this means that the eemian deposits, when found in denmark in nonglacio-dislocated state, can be regarded as being in the original position related to sea level, although there might be some movements in relation to neo-tectonic activities, as mentioned earlier. in the light of the observations mentioned above, the seven regions will be discussed according to the environmental characteristics such as the climatic affinities for the molluscs recorded in appendix 6 for each region, as appearing in fig. 103. however, for the holocene still as many as 130 species including the recent ones (95) not found as subfossil have not been dated to give their first appearance, see fig. 102: unknown arrival in holocene. at the end of each of the seven stages the molluscan communities sensu c.g.j. petersen will be presented in tables 3–9. geus bulletin no 3.pmd 09-07-2004, 09:10151 geological survey of denmark and greenland. bulletin 10, 29-32 the sedimentation and basin evolution of the kangerlussuaq basin, southern east greenland has gained renewed interest with the licensing rounds offshore the faroe islands in 2000 and 2005, as it forms an important analogy to the faroese geological setting. the faroes frontier area is in part covered by basalts and is a high-risk area with poorly known plays and sedimentary basins. it is therefore essential to obtain as much information as possible on the evolution of sedimentary basins on the rifted volcanic margins closest to the faroese islands margin. plate reconstructions of the north atlantic region indicate the former close proximity of east greenland to the faroe islands region (fig. 1), and the kangerlussuaq basin thus constitutes the most important field analogue with respect to stratigraphy, major unconformities and basin evolution. the study of the sedimentary succession in the kangerlussuaq basin, and the provenance of the sandstones in particular, will provide constraints on exploration models and may help to predict the distribution of potential reservoir sandstones in the faroese offshore basins, and eventually lead to development of play types that are new to this frontier region. this paper presents the main conclusions from two research projects: stratigraphy of the pre-basaltic sedimentary succession of the kangerlussuaq basin -volcanic basin of the north atlantic and an innovative sedimentary provenance analysis, jointly undertaken by the geological survey of denmark and greenland (geus) and casp (formerly cambridge arctic shelf programme). both projects were initiated in october 2002 and concluded in september 2005. they form part of future exploration issues programme of the faroese continental shelf (sindri programme), established by the faroese ministry of petroleum and financed by the partners of the sindri group (see acknowledgements). cretaceous–palaeogene stratigraphy field work in the 1990s and results from the present projects require revision of the existing lithostratigraphic schemes for the pre-basaltic succession in the kangerlussuaq basin. several hitherto unknown units have been discovered, and a new detailed biostratigraphy has been established incorporating east greenland and faroe–shetland sediment provenance and palaeogene sand dispersal systems michael larsen, christian knudsen, dirk frei, martina frei,thomas rasmussen and andrew g.whitham © geus, 2006. geological survey of denmark and greenland bulletin 10, 29–32. available at: www.geus.dk/publications/bull 29 ini tia l li ne of op en in g far oes uk gø01 02 155 01 mil eps 50 km greenland n orwayuk greenland shetland islands kangerlussuaq basin faroe islands orkney islands orkneyshetland platform fa ro esh et lan d bas in kangerlussuaq blosse vill e kyst edge of palaeogene basalts 6004/12-1 6004/17-1 6004/16-1z lopra-1a 206/01-1a 204/24a-7 204/19-3a 214/19-1 205/9-1 palaeogene basalts greenland geology cretaceous–palaeogene sediments crystalline basement ice international boundary sediment input well with provenance study faroese exploration wells fig. 1. map showing the location of the kangerlussuaq basin relative to the faroe islands and the faroe–shetland basin before the opening of the north atlantic. the provenance study was carried out on samples from the kangerlussuaq basin, east greenland and from the uk exploration wells marked on the map. arrows indicate the three characteristic sediment provenance areas identified in the provenance study. also shown are the three exploration wells drilled in faroese waters. age revisions based on both macrofossils and palynomorphs (fig. 2; larsen et al. 1999, 2001). the new biostratigraphy allows close correlation with the sub-basaltic successions in the offshore basins west of shetland and in the faroe islands region. most striking is the presence in both the kangerlussuaq basin and the faroe–shetland basin of a lower cretaceous shallow marine sandstone unit as well as a number of thick paleocene units dominated by sandstones. for the provenance study six sedimentary units were established (fig. 2). the first four of these are cretaceous; the palaeogene units 5 and 6 are discussed below. in the kangerlussuaq basin several regional subaerial unconformities representing sequence boundaries indicate times of sediment erosion and bypass. based on field work a depositional model for the kangerlussuaq basin has been established and prediction of reservoir sandstones in offshore basins can be carried out. palaeogene sandstones of the kangerlussuaq basin the early palaeogene succession in the kangerlussuaq basin consists of three distinct sandy units labelled 5a, 5b and 6 (fig. 2), which represent different depositional environments. graded thin-bedded sandstones dominate the lowest unit (5a); the sandstones are laterally consistent and may form amalgamated units up to 20 m thick. they were deposited in a deep marine environment and represent selandian turbidite channel-fill and distal submarine fan lobes (larsen et al. 1999). the next unit (5b) consists of well-sorted fineto medium-grained sandstones showing hummocky cross-stratification and wave ripples. the sandstones are bioturbated showing traces of ophimorpha isp. and thallasoinoides isp. the unit forms an overall coarsening upward succession and is interpreted as deposited by shallow marine sand-bars and deltaic mouth-bars. a major erosional unconformity separates the second unit from coarse-grained and pebbly sandstones of the overlying unit 6. the sandstones show large-scale trough cross-bedding and contain fragments of coal and coalified wood. unit 6 represents fluvial sandstones deposited in a proximal braided river system (fig. 3). sandstone provenance provenance studies of sediments in the faroe–shetland basin, on the orkney–shetland platform and in the kangerlussuaq basin, east greenland have been carried out using both analysis of bulk geochemical composition of the sediments and age determinations of detrital zircons. bulk geochemical analysis was carried out at geus on a total of 440 samples. of these, 171 samples were collected from the cretaceous, paleocene and eocene sedimentary rocks in the kangerlussuaq basin, while 269 represent paleocene and eocene sediments selected from cutting samples from wells in the faroe–shetland basin. in each sample, major elements were analysed by fusion wavelength dispersive x-ray fluorescence (wd-xrf) and 33 trace elements by fusion inductively coupled plasma mass spectroscopy (icpms). the analytical results provide distinctive geochemical signatures for the various units of the kangerlussuaq basin and the faroe–shetland basin. thus, selandian marine turbiditic sandstones in the kangerlussuaq basin can be shown to be derived from beach placer deposits; their high zr contents and characteristic u/ree ratios are indicative of a high heavy mineral content. this is in contrast to marine selan30 ypr tha se dan maa cam san con tur cen alb apt e ar ly l at e p al eo ce n e e o ce n e p al ae o ge n e v ic to ry f m sh et la n d g ro u p fa ro e g ro u p 50 60 70 80 90 100 110 120 ma kangerlussuaq basin faroe–shetland basin marine mudstone marine sandstone hiatus basalt fluvial sandstone 6 4 3 2 1 5b 5a fig. 2. revised stratigraphy of the pre-basaltic succession in the kangerlussuaq basin. a formal lithostratigraphic scheme including names for new formations and members is in preparation. a composite stratigraphy of the faroe–shetland basin is shown for comparison (from grant et al. 1999; ellis et al. 2002). numbers 1–6 refer to units samples for detrital zircon studies (see fig. 4). dian sediments of the faroe–shetland basin, where comparable zr and u/ree patterns are lacking, suggesting that the selandian marine sandstones in the kangerlussuaq basin and the faroe–shetland basin have different sources. age determinations of detrital zircon during the present study a total of 4347 detrital zircon grains from 47 samples were dated. the samples were from three locations: 25 from the kangerlussuaq basin, 20 from the faroe–shetland basin and two from outcrops of old red sandstone on the orkney islands. the analyses were carried out at geus using a laser ablation inductively coupled plasma mass spectrometer (la-icp-ms, quadrupole), with in situ determination of the 207pb/206pb isotopic ratios. frei et al. (2006, this volume) have compared ages obtained by u-pb dating using both conventional shrimp and high resolution sector field la-icp-ms (hr-sf-la-icp-ms) methods and demonstrated that the pb/pb dating based on 207pb/ 206pb isotopic ratios determined by la-icp-ms is a reliable method for determining detrital zircon age distributions. the observed distribution of zircon ages clearly shows that there are two distinct sources of detrital grains in the kangerlussuaq basin (fig. 1). the lower cretaceous sedimentary rocks in the basin (units 1–3, figs 2, 4) contain dominantly neoarchaean and mesoarchaean detrital zircons. the upper cretaceous sediments, however, contain a significant proportion of proterozoic detrital zircons (unit 4, figs 2, 4). the paleocene marine sandstones are characterised by a large proportion of archaean grains and variable proportions of proterozoic zircons (unit 5, figs 2, 4). the fluvial deposits of unit 6 (figs 2, 4) contain mainly archaean detrital zircons and very few proterozoic detrital zircons. known possible sources of archaean zircons are the archaean basement gneisses south and south-west of kangerlussuaq. the most probable sources of proterozoic zircons are to the north, where the caledonian orogenic belt exposes extensive areas of proterozoic granites and gneisses, as well as archaean gneiss complexes. it can be argued that units 1–3 and 6 received their main sediment input from the south, unit 4 from the north, while unit 5 may have had input from both north and south. the samples analysed from the faroe–shetland basin range in age from early paleocene to early eocene. the detrital zircon age distributions are similar in all analysed samples (not illustrated here), and include both proterozoic and archaean zircons. the proterozoic zircons exhibit a wide age distribution from neoto palaeoproterozoic, whereas the archaean zircons all fall in a narrow cluster from 2550 to 2950 ma, with a peak around 2800 ma. the almost complete absence of mesoarchaean zircons compared to the 31 fig. 4. age distribution of detrital zircons from six units of the kangerlussuaq basin. n, number of analyses; n, sample frequency. note vertical scale varies from unit to unit. for stratigraphic position of samples, see fig. 2. fig. 3. the paleocene of the kangerlussuaq area is characterised by siliciclastic sandstone units deposited by prograding shallow marine and deltaic systems. the photograph shows fine-grained, pale sandstones (unit 6, fig. 2) overlain by dark brown volcanic deposits (tuffs and lavas). person (lower left) for scale. 6: thanetian fluvial sandstone 5: selandian marine sandstone 4: maastrichtian marine mudstone 3: albian–coniacian marine mudstone 2: albian marine sandstone 1: aptian fluvial sandstone 500 1000 1500 2000 2500 3000 3500 4000 age (ma) palaeoneomesopalaeoeo-neomesoprotorozoic archaean n=105 n=100 n=100 n=100 n=100 n=100 0 4 8 12 16 0 4 8 12 0 4 8 12 16 0 1 2 3 4 5 0 4 8 12 0 4 8 12 16 n n n n n n 32 kangerlussuaq basin suggests that the sandstones from the kangerlussuaq basin and the wells we have studied in the faroe–shetland basin were sourced from different areas. greenland sands in offshore basins? it was noted by larsen et al. (1999) that the cretaceous – early palaeogene evolution of the kangerlussuaq basin is similar to the eastern margin of the faroe–shetland basin (fig. 2). ziska & andersen (2005) stressed the apparent symmetrical development of the faroe–shetland basin and thus the possibility of locating reservoir units below the basalts on the faroe islands side of the basin, similar to those already identified on the united kingdom side. sandstone provenance studies provide a possible means of testing the hypothesis that the kangerlussuaq region provided sediment to the south-western end of the faroe–shetland basin. the only paleocene sand units in the basin so far recognised as having a western source are found in well 205/9-1 (lamers & carmichael 1999), but these units do not have heavy mineral characteristics compatible with an origin from the kangerlussuaq region (whitham et al. 2004). recent studies by jolley et al. (2005), based on pollen frequency data, suggest that sediments derived from greenland are present in several wells of the faroe–shetland basin. sediment input from greenland is related to periods of low relative sea level in the palaeogene and appear to be focused along sediment transfer zones, e.g. the westray and judd transfer zones (larsen & whitham 2005). the provenance study of sandstones from the kangerlussuaq basin demonstrates that sandstones and mudstones show stratigraphic variations that are related to major sequence boundaries. furthermore, these sands show distinctive characteristics, particularly in the age distribution of detrital zircons, suggesting that they may be differentiated from sands derived from the north-west european margin (whitham et al. 2004). further studies from wells on the western part of the faroe–shetland basin and newer, not yet released data from wells drilled in faroese waters may provide the final evidence for the palaeogene sand dispersal systems discussed here. acknowledgements the companies of the sindri group are gratefully acknowledged for their financial support of the two projects and field work in the kangerlussuaq basin, southern east greenland in 2004. the sindri group comprises agip denmark bv, amerada hess (faroes) ltd., anadarko faroes company, p/f atlantic petroleum, bp amoco exploration faroes ltd., british gas international bv, dong føroyar p/f, enterprise oil exploration ltd., føroya kolvetni p/f, petro-canada faroes gmbh, phillips petroleum europe exploration ltd., shell (uk) ltd. and statoil færøyene as. references ellis, d., bell, b.r., jolley, d.w. & o’callaghan, m. 2002: the stratigraphy, environment of eruption and age of the faroes lava group, ne atlantic: in: jolley, d.w. & bell, b.r. (eds): the north atlantic igneous province. stratigraphy, tectonic, volcanic and magmatic processes. geological society (london) special publications 197, 253–269. frei, d., hollis, j.a., gerdes, a., harlov, d., karlsson, c., vasquez, p., franz, g., johansson, l. & knudsen, c. 2006: advanced in situ geochronological and trace element microanalysis by laser ablation techniques. geological survey of denmark and greenland bulletin 10, 25–28. grant, n., bouma, n. & mcintyre, a. 1999: the turonian play in the faeroe–shetland basin. in: fleet, a.j. & boldy, s.a.r. (eds): petroleum geology of northwest europe: proceedings of the 5th conference, 661–673. london: geological society. jolley, d.w., morton, a. & prince, i. 2005: volcanogenic impact on phytogeography and sediment dispersal patterns in the northeast atlantic. in: doré, a.g. & vining, b. (eds): petroleum geology: north-west europe and global perspectives – proceedings of the 6th petroleum geology conference, 969–975. london: geological society. lamers, e. & carmichael, s.m.m. 1999: the paleocene deepwater sandstone play west of shetland. in: fleet, a.j. & boldy, s.a.r. (eds): petroleum geology of northwest europe: proceedings of the 5th conference, 645–659. london: geological society. larsen, m. & whitham, a.g. 2005: evidence for a major sediment input point into the faroe–shetland basin from the kangerlussuaq region of southern east greenland. in: doré, a. & vining, b. (eds): petroleum geology: north-west europe and global perspectives – proceedings of the 6th petroleum geology conference, 913–922. london: geological society. larsen, m., hamberg, l., olaussen, s., nørgaard-pedersen, n. & stemmerik, l. 1999: basin evolution in southern east greenland: an outcrop analog for the cretaceous–paleogene basins on the north atlantic volcanic margin. american association of petroleum geologists bulletin 83, 1236–1261. larsen, m., bjerager, m., nedkvitne, t., olaussen, s. & preuss, t. 2001: pre-basaltic sediments (aptian–paleocene) of the kangerlussuaq basin, southern east greenland. geology of greenland survey bulletin 189, 99–106. whitham, a.g., morton, a.c. & fanning, c.m. 2004: insights into cretaceous–paleocene sediment transport paths and basin evolution in the north atlantic from a heavy mineral study of sandstones from southern east greenland. petroleum geoscience 10, 61–72. ziska, h. & andersen, c. 2005: exploration opportunities in the faroe islands. in: ziska, h., warming, t. & bloch, d. (eds): faroe islands exploration conference: proceedings of the 1st conference. annales societatis scientiarum færoensis supplementum 43, 146–162. authors’ address m.l.,geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. present address: dong energy, agern alle 24–26, dk-2970 hørsholm, denmark. c.k., d.f., m.f. & t.r., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ckn@geus.dk a.g.w., casp, department of earth sciences, university of cambridge, west building, 181a huntingdon road, cambridge, cb3 odh, uk. geological survey of denmark and greenland bulletin 7, 2004, p 21-24 intense drilling activity following the discovery of the siri field in 1995 has resulted in an improved understanding of the siliciclastic palaeogene succession in the danish north sea sector (fig. 1). many of the new wells were drilled in the search for oil reservoirs in sand bodies of paleocene–eocene age. the existing lithostratigraphy was based on data from a generation of wells that were drilled with deeper stratigraphic targets, with little or no interest in the overlying palaeogene sediments, and thus did not adequately consider the significance of the palaeogene sandstone units in the danish sector. in order to improve the understanding of the distribution, morphology and age of the palaeogene sediments, in particular the economically important sandstone bodies, a detailed study of this succession in the danish north sea has recently been undertaken. an important aim of the project was to update the lithostratigraphic framework on the basis of the new data. the project was carried out at the geological survey of denmark and greenland (geus) with participants from the university of aarhus, dong e&p and statoil norway, and was supported by the danish energy agency. most scientific results cannot be released until september 2006, but a revised lithostratigraphic scheme may be published prior to that date. formal definition of new units and revision of the lithostratigraphy are in preparation. all of the widespread palaeogene mudstone units in the north sea have previously been formally established in norwegian or british wells, and no reference sections exist in the danish sector. as the lithology of a stratigraphic unit may vary slightly from one area to another, danish reference wells have been identified during the present project, and the lithological descriptions of the formations have been expanded to include the appearance of the units in the danish sector. 21 a revised lithostratigraphy for the palaeogene – lower neogene of the danish north sea poul schiøler, jan andsbjerg, ole r. clausen, gregers dam, karen dybkjær, lars hamberg, claus heilmann-clausen, lars e. kristensen, iain prince and jan a. rasmussen geological survey of denmark and greenland bulletin 7, 21–24 (2005) © geus, 2005 5°e 4°e 57°n 56°n c offee so il fault 50 km 20 40 60 80 100 thickness (m) siri-3 siri canyon f-1x inez-1 siri-1 siri-2 c entral g raben north sea denmark fig. 1. map showing the study well data base superposed on an isochore map of the lista formation. the lista formation thickens towards the northern part of the mapped area; this thickening is mainly due to the presence of sandstone bodies within the siri canyon. many of the sandstone bodies recently discovered in the danish sector have a limited spatial distribution and were sourced from other areas than their contemporaneous counterparts in the norwegian and british sectors. these sandstone bodies are therefore defined as new lithostratigraphic units in the danish sector, and are assigned danish type and reference sections. there is a high degree of lithological similarity between the palaeogene–neogene mudstone succession from danish offshore boreholes and that from onshore exposures and boreholes, and some of the mudstone units indeed seem identical. however, in order to acknowledge the traditional distinction between offshore and onshore stratigraphic nomenclature, the two sets of nomenclature are kept separate herein. in recent years oil companies operating in the north sea have developed various in-house lithostratigraphic charts for the paleocene–eocene sand and mudstone successions in the danish and norwegian sectors. a number of informal lithostratigraphic units have been adopted and widely used. in the present project, these units have been formally defined and described, maintaining their original names whenever feasible, with the aim of providing an unequivocal nomenclature for the palaeogene – lower neogene succession in the danish sector. it has not been the intention to establish a sequence stratigraphic model for this succession in the north sea; the reader is referred to the comprehensive works of michelsen (1993), neal et al. (1994), mudge & bujak (1994, 1996a, b), michelsen et al. (1995, 1998), danielsen et al. (1997) and rasmussen (2004). methodology the multidisciplinary project integrates biostratigraphic results from palynology and micropalaeontology with results from sedimentological studies, petrophysical log interpretation and conventional seismic interpretation, in an attempt to identify, describe and correlate lithological units. the lithostratigraphy proposed is based on analysis of petrophysical logs from more than 80 wells (fig. 1) and studies of cuttings samples and core sections. lithostratigraphic well correlation is supported by results from biostratigraphic analysis of 29 wells encompassing both palynological and micropalaeontological data. this analysis has established a detailed succession of first downhole occurrences of biostratigraphical events which are of crucial importance for well correlation and age determination in the region. resulting well correlations have been matched with results from interpretation of key seismic sections. a revised lithostratigraphy an important outcome of the project is a revised lithostratigraphic chart for the siliciclastic palaeogene – lower neogene succession of the danish north sea (fig. 2). the chart is based on the subdivision proposed by deegan & scull (1977), but incorporates subsequent improvements by hardt et al. (1989) and knox & holloway (1992). the revised lithostratigraphy presented in fig. 2 has its genetic base at the top of the chalk group (ekofisk for22 fig. 2. diagram showing the correlation between key lithostratigraphic schemes for the central graben and eastern north sea. white: mudstone units. yellow: sandstone units. names in bold indicate units of formation or group rank. nordland group hordaland group balder sele lista unnamed unit ekofisk cen5 cen4 cen3 cen2 cen1 north sea marl ekofisk nordland group hordaland group balder sele lista våle ekofisk nordland group lark horda balder taym o u sa h ef ri n g d u fa fr ej a k o lg a f u r r in d id u n ty r b o r sele lista l is ta fo rt ie s c ro m ar ty m ey v a d e f o rt ie s a n d re w f is k e b a n k f is k e b a n k f ri g g maureen m a u re e n ekofisk ekofisk nordland group lark horda balder sele bue ve vile våle deegan & scull (1977) kristoffersen & bang (1982) hardt et al. (1989) knox & holloway (1992) this study r o g a la n d g ro u p mation). the top is at the mid-miocene unconformity, a basinwide erosion surface that separates the oligocene to middle miocene hordaland group from the overlying middle miocene to recent nordland group. the succession is subdivided into seven formations containing eleven new members. the våle, lista, sele, fur, balder, horda and lark formations of previous lithostratigraphic schemes are adequate for subdividing the danish sector at formation level and are retained herein with expanded lithological descriptions. the five first-mentioned formations constitute the rogaland group of early selandian (late paleocene) to earliest eocene (pre-classical ypresian) age, and the horda and lark formations constitute the hordaland group of ypresian (early eocene) to early serravallian (middle miocene) age. rogaland group the våle formation (hardt et al. 1989) consists of light grey to greenish grey marlstones and is retained largely unchanged from its original description. the lista formation (deegan & scull 1977) is subdivided into three new mudstone members (vile, ve and bue). the predominantly light to dark grey mudstones of the vile and bue members are separated by the greenish, bluish to red coloured mudstones of the ve member. the sele formation (deegan & scull 1977) overlies the lista formation; the boundary is placed where laminated dark grey to black mudstones overlie the light to dark grey, smectitic, non-laminated mudstones of the bue member. this boundary definition differs from that used by knox & holloway (1992), who placed the lista–sele boundary at the base of the bue member as recognised here (fig. 2). incidentally, the lista–sele boundary as originally defined coincides with the paleocene–eocene boundary recently approved by the international commission of stratigraphy, at the conspicuous 13c isotope excursion in chron c24r. the balder formation is retained largely unchanged from its original description by deegan & scull (1977). the fur formation (pedersen & surlyk 1983) is characterised by interbedded diatomite and tuffs and has been recognised in one norwegian and three danish offshore wells by thomsen & danielsen (1995) at a similar stratigraphic level to that of the sele and balder formations. it is therefore included in the north sea lithostratigraphic chart (fig. 2). hordaland group the horda and lark formations of knox & holloway (1992) are introduced in the danish sector for the basinal mudstones that overlie the balder formation. the horda formation consists of greenish-grey to greyish-green fissile mudstone, whereas the lark formation comprises greenishgrey to dark brown non-fissile mudstones. the transition from the horda to the lark formation is characterised by a significant increase in gamma ray response. sandstone members eight discrete sandstone members have been recognised in the palaeogene – lower neogene mudstone succession (fig. 2). the lower five of these occur exclusively in the siri canyon, a 20 km wide and up to 200 m deep erosional depression that truncates reflectors in the chalk group (fig. 1; hamberg et al. 2005). the bor member is a sandstone unit in the våle formation. three sandstone members occur in the lista formation and are named tyr, idun and rind. the sele and horda formations each contain sandstone units for which the names kolga and hefring members are proposed, respectively. the hefring member occurs in an area south of the siri canyon. two sandstone members (dufa and freja) are present in the lark formation. the dufa member occurs in the north-eastern part of the danish sector, around the wells f-1x and inez-1 (fig. 1); the freja member occurs above the siri canyon, south of the siri-1 well (fig. 1). rogaland group sandstones: reservoir facies the economically important paleocene–eocene oil reservoir sandstones of the rogaland group are situated at the mouth of, and within, the siri canyon (fig. 1). the sandstones occur as in situ sediment bodies deposited by gravity flows and as remobilised, injected sandstones (hamberg et al. 2005). although the sandstone bodies occur at different stratigraphic levels within the rogaland group, they are strikingly uniform. the sandstones are quartzose and very fineto finegrained, well sorted and rich in glaucony with a greenish grey colour. the sandstone bodies appear massive and structureless with primary sedimentary structures preserved only in their basal part. individual sandstones may be up to 100 m thick but are usually much thinner. identification and correlation of individual sandstone members are feasible by means of 3d seismic data, with support from borehole logs and high-resolution biostratigraphy. an example of results from such an integrated correlation in the lower part of the siri canyon is shown in fig. 3. 23 acknowledgements this work was made possible through grants from the danish energy authority, under the energy research programme 2000. dong e&p is thanked for access to lithostratigraphic data from siri canyon wells. references danielsen, m., michelsen, o. & clausen, o.r. 1997: oligocene sequence stratigraphy and basin development in the danish north sea sector based on log interpretations. marine and petroleum geology 14, 931–950. deegan, c.e. & scull, b.j. (compilers) 1977: a standard lithostratigraphic nomenclature for the central and northern north sea. institute of geological sciences report 77/25; norwegian petroleum directorate bulletin 1, 36 pp. hamberg, l., dam, g., wilhelmson, c. & ottesen, t.g. 2005: paleocene deep-marine sandstone plays in the siri canyon, offshore denmark – southern norway. in: doré, a.g. & vining, b.a. (eds): petroleum geology: north west europe and global perspectives. proceedings of the 6th petroleum geology conference, 1185–1198. london: geological society. hardt, t., holtar, e., isaksen, d., kyllingstad, g., lervik, k.s., lycke, a.s. & tonstad, k. 1989: revised tertiary lithostratigraphic nomenclature for the norwegian north sea. in: isaksen, d. & tonstad, k. (eds): a revised cretaceous and tertiary lithostratigraphic nomenclature for the norwegian north sea. norwegian petroleum directorate bulletin 5, 35–55. knox, r.w.o’b. & holloway, s. 1992: paleogene of the central and northern north sea. in: knox, r.w.o’b. & cordey, w.g. (eds): lithostratigraphic nomenclature of the uk north sea, 133 pp. nottingham: british geological survey. kristoffersen, f.n. & bang, i. 1982: cenozoic excl. danian limestone. in: michelsen, o. (ed.): geology of the danish central graben. danmarks geologiske undersøgelse serie b 8, 62–71. michelsen, o. 1993: stratigraphic correlation of the danish onshore and offshore tertiary successions based on sequence stratigraphy. bulletin of the geological society of denmark 41, 145–161. michelsen, o., danielsen, m., heilmann-clausen, c., jordt, h., laursen, g. & thomsen, e. 1995: occurrence of major sequence stratigraphic boundaries in relation to basin development in cenozoic deposits of the southeastern north sea. in: steel, r.j. et al. (eds): sequence stratigraphy on the northwest european margin. norsk petroleum forening special publication 5, 415–427. michelsen, o., thomsen, e., danielsen, m., heilmann-clausen, c., jordt, h. & laursen, g. 1998: cenozoic sequence stratigraphy in the eastern north sea. in: de graciansky, p.-c. et al. (eds): mesozoic and cenozoic sequence stratigraphy of european basins. society of economic paleontologists and mineralogists special publication 60, 91–118. mudge, d.c. & bujak, j.p. 1994: eocene stratigraphy of the north sea basin. marine and petroleum geology 11, 166–181. mudge, d.c. & bujak, j.p. 1996a: paleocene biostratigraphy and sequence stratigraphy of the uk central north sea. marine and petroleum geology 13, 295–312. mudge, d.c. & bujak, j.p. 1996b: an integrated stratigraphy for the paleocene and eocene of the north sea. in: knox, r.w.o’b., corfield, r.m. & dunay, r.e. (eds): correlation of the early paleogene in northwest europe. geological society special publication (london) 101, 91–113. neal, j.e., stein, j.a. & gamber, j.h. 1994: graphic correlation and sequence stratigraphy in the palaeogene of nw europe. journal of micropalaeontology 13, 55–80. pedersen, g.k. & surlyk, f. 1983: the fur formation, a late paleocene ash-bearing diatomite from northern denmark. bulletin of the geological society of denmark 32, 43–65. rasmussen, e.s. 2004: the interplay between true eustatic sea-level changes, tectonics, and climate changes: what is the dominating factor in sequence formation of the upper oligocene – miocene succession in the eastern north sea basin, denmark? global and planetary change 41, 15–30. thomsen, e. & danielsen, m. 1995: transitional paleocene/eocene ashbearing diatomite in the eastern north sea. tertiary research 15, 111–120. siri-2 siri-1 siri-3 rind mb idun mb kolga mb tyr mb tyr mb horda fm sli dbf fan aau ama ppy balder fm sele fm lista fm våle fm 50 m ipf gda fig. 3. correlation diagram of the rogaland group in the lower part of the siri canyon (well locations shown in fig. 1). horizontal distance not to scale. in this section the tyr member consists of three sandstone beds separated by mudstone layers. the idun member cuts deeply into the underlying lista formation. key biostratigraphic downhole events are shown to the left. sli, subbotina linaperta; dbf, decrease in benthic foraminifera; fan, influx of fenestrella antiqua; aau, apectodinium augustum; ama, alisocysta margarita; ppy, common palaeoperidinium pyrophorum; ipf, increase in planktonic foraminifera; gda, globoconusa daubjergensis. authors’ addresses p.s., j.a., k.d., l.e.k. & j.a.r., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: pos@geus.dk o.r.c. & c.h.-c., department of earth sciences, university of aarhus, dk-8000 århus c, denmark. g.d. & l.h., dong, agern allé 24–26, dk-2970 hørsholm, denmark. i.p., statoil norway, forusbeen 50, n-4035 stavanger, norway. 24 geological survey of denmark and greenland bulletin 3, 24-97 24 the danish late quaternary marine molluscs the record of shell-bearing danish late quaternary marine molluscs has been established on the basis of finds made during the systematic geological mapping since 1888, as presented in the publications from the geological survey of denmark, mainly in the i. række covering the descriptions for the map sheets. furthermore, special papers on holocene and late pleistocene marine molluscs have been included. most of them have been listed in one of the preceding chapters on works on danish sites with marine sediments. 1. the bælt sea area 2. the baltic sea area 3. the kattegat area with fjords 4. the limfjord area 5. the north sea coastal area 6. the vendsyssel area 7. the skagen area the regions are figured on the map (fig. 2) and follow the outline of the geological map sheets except the skagen area. it should be especially noted that the southern limit during the holocene of the north sea coastal region is at blåvands huk. although this area from varde to the german border has been mapped, no descriptions have been published. however, studies of the eemian from this region have been published and will be commented upon when presenting the eemian records from the above-mentioned areas. the bælt sea area has been taken as the region between the kattegat region and the baltic, here following ekman (1953) saying that the boundary between the bælt sea and the baltic proper is the threshold between gedser and darss and the southern end of the øresund, see fig. 1. the reason for including the øresund north of saltholm (northern and middle øresund sensu jensen & knudsen 1995) in the kattegat region is that regarding the water from the baltic the øresund is a sound – a passage – for the brackish water flowing north, but regarding the salt-water from the kattegat, the øresund is a fjord down to the threshold between amager and limhamn, to quote thorson (1944a, p. 42). the subfossil shell-bearing molluscs are presented on the background of the annotated check list of recent marine molluscs of danish waters by jensen & knudsen (1995). furthermore, general information has been taken from jensen & spärck (1934), lemche (1948), nordsieck (1968, 1969), fretter & graham (1976–1978, 1980–1982, 1984), fretter et al. (1986) and poppe & goto (1991, 1993) (which will not be quoted throughout the text, listing the many species and their environment). only the synonyms mentioned in texts on danish molluscan finds are included. a list of these synonyms is found heading the index of species. the investigations carried out by scientists in the northern atlantic have given a good base for the evaluation of the danish late quaternary molluscan fauna. these studies have been published mainly in papers on the zoology of east greenland, the godthaab expedition 1928, the zoology of iceland, and the zoology of the faroes. however, the zoology of greenland has been supplemented by contributions from macpherson (1971), lubinsky (1980), the 6. og 7. thule expedition til sydøstgrønland 1931–33 under the leadership of knud rasmussen and the treaarsexpeditionen til christian den x’s land 1931–34 under the leadership of lauge koch. in the two last-mentioned contributions, especially the animal ecology and the arctic communities have been treated, which form a very important part in the discussion of the danish late quaternary molluscan assemblages. the following publications on molluscs can be mentioned: 1. from greenland: spärck (1933), thorson (1933, 1944b, 1951), thorson & ussing (1934), madsen (1936), bertelsen (1937), lemche (1941a, b), ockelmann (1958), kramp (1961, 1963). 2. from iceland: spärck (1937), lemche (1938), thorson (1941), knudsen (1949a, b), madsen (1949). 3. from the faeroes: lemche (1928), thorson & spärck (1928), spärck & thorson (1931), petersen (1968), knudsen (1970). as to the community concept as worked out by c.g.j. petersen, references have already been given to petersen & jensen (1911), petersen (1913, 1914, 1915, 1918), and thorson (1957). the regional division of the northern european seas in to arctic, subarctic, boreal and lusitanian is from the zoogeographical division of the geus bulletin no 3.pmd 28-06-2004, 08:4524 25 northern european seas as given by feyling-hanssen (1955, fig. 5) and símonarson et al. (1998), as seen on figs 4 and 5 respectively. the molluscan genera and species are presented in groups within class, subclass and order mainly following the presentation of recent marine molluscs of danish waters as given by jensen & knudsen (1995). when there is no subfossil record at hand the information is taken from the recent data as given by poppe & goto (1991, 1993), jensen & knudsen (1995) and others as listed previously. to facilitate the use of the index of all molluscan species, a list of synonyms is given as mentioned earlier including the species mentioned in danish mollusc literature. class polyplacophora order neoloricata among the seven species recorded in the recent danish fauna only one, tonicella marmorea, has been found subfossil in the vendsyssel area from the younger weichselian deposits. it is a circumboreal species which in europe is known from northern scandinavia south to denmark and recorded from around iceland (knudsen 1949a, b) down to ireland (poppe & goto 1991). it is common in danish waters, being known from the central part of the kattegat and the øresund region (muus 1959), e.g. subarctic–boreal–lusitanian species. habitat. common in shallow water of less than 20 m, otherwise recorded from 0–183 m (muus 1959). the lack of information on subfossil finds may reflect difficulty in determination when the species is found in the subfossil state of preservation with the shell parts apart. polyplacophora do occur as seen from the skagen well where finds have been recognised at the 70.10–70.30 m and the 37.0–37.25 m levels, viz. during the subboreal and the subatlantic respectively. among the other – only recent – finds of polyplacophora, leptochiton asellus is widely found from the arctic to the lusitanian. hanleya hanleyi, ischnochiton albus and the above-mentioned tonicella marmorea are all found in the subarctic, boreal and lusitanian regions. the last three species, callochiton septemvalvis, lepidochitona cinereus and tonicella rubra, are all restricted to the boreal–lusitanian region. class gastropoda subclass prosobranchia order archaeogastropoda scissurella crispata fleming 1828 fig. 6 distribution. w greenland, s and w iceland, spitsbergen, norway north of lofoten, and south to the mediterranean. occurrence. the subarctic (not in true arctic water), boreal and lusitanian regions (according to thorson 1941). habitat. in the waters around iceland (thorson 1941, p. 4), the living specimens have often been found at depths greater than 500 m on clay bottom. however, poppe & goto (1991, p. 64) write that the species lives on stones, shelly sand and clay bottoms between 15 and 600 m. fretter & graham (1976, pp. 2–4) stated that the species is always sublittoral, even in the extreme northerly limits of its range, and occurs from 8– 2000 m. it is not recorded from danish waters although fretter & graham mentioned it from the norwegian and swedish coasts of the skagerrak. only subfossil finds. during the eemian in the kattegat region. fig. 6. scissurella crispata fleming 1828. anholt ii, 78.06–78.09 m b.s. × 20. mguh 25316. geus bulletin no 3.pmd 28-06-2004, 08:4525 26 patella vulgata linnaeus 1758 distribution. faeroes, norway off the lofoten islands, and south to the straits of gibraltar (thorson 1941). occurrence. boreal–lusitanian. habitat. intertidal, rocky shores or man-made hard substrates. however, it also occurs with seaweed. only one recent individual has been found at løkken, while empty shells are often found on the skagerrak coast (knudsen 1993). subfossil finds. the limfjord region, holocene, in a ‘køkkenmødding’ (kitchen midden) at bulbjerg (petersen 1888). helcion pellucidum (linnaeus 1758) distribution. west and south iceland, northern norway north of lofoten, and south to portugal (thorson 1941). also found in the øresund, but absent from the baltic and the limfjord. occurrence. boreal and lusitanian. habitat. lives on seaweeds at depths from 0 to 27 m. petersen (1888) points to this habitat as a reason why it is rarely found. subfossil finds. the limfjord region and vendsyssel, holocene. lepeta caeca (müller 1776) distribution. from northern scandinavia, n and s iceland and spitsbergen (thorson 1941) south to scotland. it is not present in the baltic, the north sea and the channel, but occurs in the azores. this species, according to poppe & goto (1991), prefers cold temperatures and lives at greater depths in the southern part of its range. in this way it exemplifies the tropical submerge. it is present in the øresund region and has been reported from single finds in the kattegat (petersen 1888). occurrence. subarctic and boreal. subfossil finds. none. iothia fulva (müller 1776) distribution. northern scandinavia and s and w iceland south to the irish sea, and like lepeta caeca in deep water off the azores. it is found in the øresund region, but with rare and single finds in the kattegat (petersen 1888). occurrence. boreal and lusitanian. habitat. offshore between 5 and 600 m on hard substrates like lepeta caeca. subfossil finds. the limfjord region, holocene. acmaea tessulata (müller 1776) distribution. east greenland, around iceland (thorson 1941) and in scandinavia according to petersen (1888). the limfjord (petersen 1986a). common in the øresund, but absent in the baltic. extends southwards to the north of the british isles and northern ireland. occurrence. arctic, subarctic, boreal, and lusitanian. habitat. lower part of the intertidal zone. in the south connected with zostera. subfossil finds. the limfjord region, holocene. acmaea virginea (müller 1776) distribution. northern scandinavia and around iceland (thorson 1941), south to the cape verde islands. common in the kattegat and øresund, but not found in the baltic. occurrence. subarctic, boreal, lusitanian. habitat. on hard substrate at depths from 0 to 100 m. subfossil finds. the limfjord and vendsyssel regions, holocene. emarginula fissura (linnaeus 1758) distribution. scandinavia and south to the mediterranean, according to petersen (1888) taken alive from the øresund, dead shells in the northern kattegat. geus bulletin no 3.pmd 28-06-2004, 08:4526 27 occurrence. boreal and lusitanian. habitat. from the low tide line to a depth of 700 m on hard substrate. subfossil finds. none. puncturella noachina (linnaeus 1771) distribution. spitsbergen and around iceland (thorson 1941), scandinavia south to portugal. also found in the skagerrak and kattegat including øresund. occurrence. arctic, subarctic, boreal and lusitanian. habitat. between 10 and 200 m on rock and stones. subfossil finds. none. margarites helicinus (phipps 1774) distribution. from northern scandinavia, spitsbergen and around iceland (thorson 1951) south to the british isles. a few records from the skagerrak and kattegat. occurrence. arctic, subarctic, boreal and northern part of the lusitanian region. habitat. from the intertidal zone to 400 m deep on seaweeds and under stones. subfossil finds. the limfjord region, holocene. gibbula cineraria (linnaeus 1758) distribution. from northern scandinavia north of lofoten, and w and s iceland (thorson 1941) south to morocco. according to petersen (1888) known from the kattegat including øresund, the bælt sea and the limfjord area. occurrence. boreal–lusitanian. habitat. intertidal to 130 m deep on rocks and seaweeds. subfossil finds. the limfjord, the north sea and vendsyssel, holocene. recorded from the north sea during the eemian. gibbula tumida (montagu 1803) distribution. from northern norway north of lofoten, and s and w iceland (thorson 1941), south to spain. known from the kattegat, including the øresund, but not so common as gibbula cineraria. the species does not occur in the limfjord (petersen 1986a, table 1). occurrence. boreal–lusitanian. habitat. on gravel bottoms from below low tide to depths of 1200 m. subfossil finds. the limfjord and vendsyssel area, holocene. jujubinus clelandi (w. wood 1828) distribution. from the lofoten islands south into the mediterranean. found in the kattegat region, including the øresund, but rare. occurrence. boreal–lusitanian. habitat. on various types of bottom from depths of 35 m to 800 m. subfossil finds. none. calliostoma formosa (mighels 1842) distribution. along the norwegian coast, w and s iceland (thorson 1941), and south to the british isles but not on the west side. occurrence. boreal. habitat. dredged at depths between 19 and 1000 m – the species is never littoral. subfossil finds. none. calliostoma zizyphinum (linnaeus 1758) distribution. from the lofoten islands south to the azores. in the skagerrak. occurrence. boreal–lusitanian. geus bulletin no 3.pmd 28-06-2004, 08:4527 28 habitat. intertidal to 300 m deep – lives on all types of bottoms. subfossil finds. none. skenea serpuloides (montagu 1808) distribution. from the british isles south to portugal and into the mediterranean. occurrence. lusitanian. habitat. from the intertidal zone down to 50 m deep. intertidal on weeds and stones, and sublittoral dredged from shelly and gravelly sand. only subfossil finds. the limfjord area, holocene. skenea basistriata (jeffreys 1877) distribution. atlantic coast of europe, but not in the north sea and the baltic. occurrence. boreal–lusitanian. habitat. on soft bottom – never in shallow water or littoral sequences, deep water 90–2400 m. subfossil finds. the limfjord region, holocene. theodoxus fluviatilis (linnaeus 1758) distribution. from the pyrenees and the british isles east towards the caucasus, including northern sweden and the coasts of finland. occurrence. lusitanian–boreal. habitat. the primary habitat of this species is rivers. in the baltic sea the form littoralis becomes a common littoral animal (fretter & graham 1978a, p. 105). the species is also recorded from the fjords bordering the kattegat region today. subfossil finds. the kattegat and limfjord regions, holocene. the archaeogastropoda are represented by 16 species in the recent marine fauna, out of which ten have been recorded from the past. only two species, skenea serpuloides and scissurella crispata, are not found in the recent fauna. there is a total number of subfossil finds among the archaeogastropoda of 12 species, with two from the eemian and ten from the holocene. order mesogastropoda littorina littorea (linnaeus 1758) distribution. from northern norway north of lofoten, and south to spain. common along all the danish coasts but not in the baltic and on the more exposed sandy coasts (petersen 1888). recent records from the baltic as far as bornholm (fretter & graham 1980, p. 256). occurrence. boreal and lusitanian. habitat. the intertidal zone, abundant on rocky shores, might be found to a depth of 60 m. subfossil finds. the bælt sea, baltic, kattegat, limfjord, the north sea and vendsyssel regions, holocene. the bælt sea, baltic, kattegat and north sea regions in the eemian. melaraphe (littorina) neritoides (linnaeus 1758) distribution. from western norway south to morocco and the mediterranean. a scattered occurrence in the kattegat (jensen & knudsen 1995). the species is recorded from the limfjord (petersen 1986a). occurrence. boreal and lusitanian. habitat. lives high on the rocky shores (splash zone). subfossil finds. none. littorina mariae sacchi & rastelli 1966 distribution. from northern scandinavia south to the mediterranean, extending through the kattegat into the bælt sea. occurrence. the boreal and lusitanian. habitat. in the tidal zone on weeds. subfossil finds. none. (the species may have been confused with littorina obtusata.) geus bulletin no 3.pmd 28-06-2004, 08:4528 29 littorina obtusata (linnaeus 1758) distribution. w greenland, s and w iceland (thorson 1941), northern scandinavia north of lofoten, and south to the mediterranean. the species extends through the limfjord and kattegat into the bælt sea (fretter & graham 1980). occurrence. subarctic, boreal and lusitanian. habitat. intertidal, lives on weeds. subfossil finds. the bælt sea, kattegat, limfjord, north sea and vendsyssel regions, holocene. during the eemian recorded from the north sea. littorina saxatilis (olivi 1792) distribution. from greenland, spitsbergen, around iceland and the atlantic coasts of europe. common in the fjords bordering the kattegat, including the limfjord (petersen 1986a). occurrence. the arctic, subarctic, boreal and lusitanian. habitat. intertidal. subfossil finds. the bælt sea, kattegat, limfjord, north sea and vendsyssel regions, holocene. the north sea during the eemian, and the vendsyssel area in the late weichselian. littorina tenebrosa (montagu 1803) distribution. off southern iceland and the atlantic coasts of europe. the species is found in the limfjord, and penetrates also into the baltic, with finds off møn and stevns (petersen 1888). occurrence. the boreal and lusitanian. habitat. intertidal. subfossil finds. the bælt sea, baltic, kattegat, limfjord and vendsyssel regions, holocene. lacuna pallidula (da costa 1778) fig. 7a, b distribution. from spitsbergen, around iceland and along the atlantic coast down to the gulf of biscay. enters the danish waters, including the limfjord, the øresund and the bælt sea. occurrence. the arctic, subarctic, boreal and lusitanian. habitat. intertidal to 70 m deep on weeds. subfossil finds. the bælt sea, limfjord, north sea and vendsyssel regions, holocene. in the skagen well from the subatlantic. lacuna crassior (montagu 1803) distribution. from the arctic seas to the british isles. only records from the north sea and from the nw coast of sweden. occurrence. the arctic, subarctic and boreal. habitat. sublittoral to 90 m deep on soft bottoms with stones and shells (fretter & graham 1980, p. 248). subfossil finds. none. fig. 7. a, b: lacuna pallidula (da costa 1778). skagen 4, 30.8– 30.5 m b.s., lab. no. 355,93. × 20. mguh 25317. geus bulletin no 3.pmd 28-06-2004, 08:4529 30 lacuna parva (montagu 1803) distribution. from norway off the lofoten, and south to spain, found in the kattegat region with fjords, but few records, and not found in the limfjord. occurrence. boreal and lusitanian. habitat. intertidal extending sublittorally to around 50 m as lacuna vincta (fretter & graham 1980, p. 250) and living on seaweeds. subfossil finds. the limfjord, north sea and vendsyssel regions, holocene. from the north sea during the eemian. lacuna vincta (montagu 1803) distribution. from w greenland around iceland, norway and south to spain. in the danish waters, including the limfjord, found into the bælt sea (fretter & graham 1980). occurrence. subarctic, boreal and lusitanian. habitat. intertidal to 60 m deep living on seaweeds. subfossil finds. the bælt sea, kattegat, limfjord, north sea and vendsyssel regions, holocene. in the north sea region found during the eemian. in the vendsyssel area recorded from the eemian, and also from the early/middle and late weichselian (older and younger yoldia sea respectively). hydrobia neglecta muus 1963 distribution. the british isles, ireland and the north sea. occurrence. boreal and lusitanian. habitat. shallow-water environments on the soft substratum or the vegetation. subfossil finds. none. hydrobia ulvae (pennant 1777) fig. 8 distribution. norway off lofoten south to the mediterranean. in all the danish waters including the baltic. occurrence. the boreal and lusitanian. habitat. the intertidal zone, but has been found as deep as 20 m, on soft substrate, most often on intertidal banks of firm mud or muddy sand (fretter & graham 1978a, p. 122). subfossil finds. the bælt sea, baltic, kattegat, limfjord, north sea and vendsyssel regions, holocene. in the skagen well from the subatlantic. during the eemian recorded from the bælt sea, baltic, kattegat and the north sea regions. hydrobia ventrosa (montagu 1803) distribution. norway off lofoten, south to the mediterranean. in all the danish waters including the baltic. occurrence. the boreal and lusitanian regions. habitat. the intertidal zone,on soft substratum like hydrobia ulvae, but prefers lower salinities (fretter & graham 1978a, p. 126). the two species may occur together, so quantitative analyses must be undertaken to designate any changes in the environment (petersen 1993). subfossil finds. the baltic, kattegat, limfjord and the north sea regions, holocene. potamopyrgus antipodarum (gray 1853) distribution. from scandinavia to spain and in danish waters into the baltic, a late immigrant according to jensen & knudsen (1995). occurrence. the boreal and lusitanian regions. fig. 8. hydrobia ulvae (pennant 1777). skagen 4, 25.0–25.5 m b.s., lab. no. 350,93. × 20. mguh 25318. geus bulletin no 3.pmd 28-06-2004, 08:4530 31 habitats. in all kinds of brackish and freshwater habitats. much like the distribution of hydrobia ventrosa in brackish waters (fretter & graham 1978a, p. 132). subfossil finds. none. skeneopsis planorbis (fabricius 1780) distribution. w greenland, around iceland and norway, south to the mediterranean. known from few places in danish waters. occurrence. the subarctic, boreal and lusitanian regions. habitat. intertidal down to a depth of 70 m, lives on seaweeds. only subfossil finds. the limfjord and the vendsyssel regions, holocene. barleeia unifasciata (montagu 1803) fig. 9 distribution. from the shetlands south into the mediterranean. occurrence. the boreal and lusitanian regions. habitat. shallow waters, intertidal, lives on seaweeds on rocky shores. only subfossil finds. the skagen well from the subatlantic. alvania abyssicola (forbes 1850) distribution. from northern norway to the mediterranean. the species extends into the skagerrak and kattegat, including the øresund. occurrence. the boreal and lusitanian regions. habitat. on muddy bottom in sublittoral areas at depths of 15–100 m. subfossil finds. the vendsyssel area from the eemian. alvania jeffreysi (waller 1864) distribution. s and w iceland, norway, south to the mediterranean. found in the skagerrak, but not in the north sea. occurrence. the boreal and lusitanian regions. habitat. always sublittoral from 50 to 600 m on sandy bottom. subfossil finds. none. alvania lactea (michaud 1830) distribution. from the channel islands south to morocco and the mediterranean. occurrence. the lusitanian region. habitat. sublittorally under stones and amongst algae. only subfossil finds. the limfjord and vendsyssel areas, holocene. alvania cimicoides (forbes 1844) distribution. sw and nw iceland (empty shells), norway, north of lofoten, and south to the mediterranean; probably absent from the channel and the north sea. occurrence. the boreal and lusitanian regions. habitat. sublittoral, from the laminarian zone downwards, but mainly in deeper water down to 500 m. fig. 9. barleeia unifasciata (montagu 1803). skagen 3, 33.90– 34.20 m b.s., lab. no. 709,93. × 20. mguh 25319. geus bulletin no 3.pmd 28-06-2004, 08:4531 32 usually found on soft bottoms (fretter & graham 1978b). only subfossil finds. the vendsyssel area, holocene. alvania punctura (montagu 1803) distribution. norway south of lofoten (including lofoten) and south to the mediterranean. it extends to the swedish west coast, but is absent from the øresund, the baltic, the eastern shores of the north sea and the eastern basin of the channel (fretter & graham 1978b). however, according to jensen & knudsen (1995) alvania punctura is found in the northern and central parts of the øresund. occurrence. the boreal and lusitanian regions. habitat. sublittoral to depths of c. 100 m, on both finer and coarser substrata. subfossil finds. the limfjord and the vendsyssel regions, holocene. alvania cruenta odhner 1915 distribution. arctic canada, west greenland and svalbard (thorson 1951; macpherson 1971). occurrence. the arctic and subarctic. habitat. from 19 to 234–254 m on mud (macpherson 1971). only subfossil finds. the early/middle weichselian in the vendsyssel region. alvania jan mayeni (friele 1886) distribution. e and w greenland, spitsbergen, ne iceland, and norway north of lofoten (thorson 1941). occurrence. the arctic and subarctic with boreal outposts. habitat. off iceland between 94–442 m in deep on clay with many stones (thorson 1941). only subfossil finds. the early/middle weichselian (older yoldia clay) in vendsyssel. alvania scrobiculata (möller 1842) distribution. e and w greenland, spitsbergen, n and e iceland, and norway north of lofoten (thorson 1941). occurrence. the arctic and subarctic with boreal outposts. habitat. from 22 m at e greenland to 342 m in the northern arctic sea on a bottom of sand and with algae (thorson 1941). only subfossil finds. from the early/middle weichselian (older yoldia clay) in vendsyssel. cingula semistriata (montagu 1808) distribution. from lofoten and south along the west coast of norway to the mediterranean (rare in the north). it extends into the kattegat, but is absent from the eastern shores of the southern north sea and the limfjord. occurrence. the boreal and lusitanian regions. habitat. on rocky shores in the intertidal zone and sublittorally to 100 m, fond of silty places (fretter & graham 1978b). subfossil finds. the limfjord and the vendsyssel areas, holocene. cingula turgida (jeffreys 1870) distribution. from norway north of lofoten to the south into kattegat. occurrence. the boreal region. habitat. on muddy bottoms down to c. 1000 m. subfossil finds. the north sea region, holocene. obtusella alderi (jeffreys 1858) distribution. from norway to spain, not on the eastern shores of the north sea and in the baltic; however, found in the kattegat, including the northern part of the øresund. occurrence. the boreal and lusitanian regions. geus bulletin no 3.pmd 28-06-2004, 08:4532 33 habitat. sublittoral to a depth of 60 m amongst algae and on sandy or gravelly bottoms. subfossil finds. none. onoba aculeus (gould 1841) distribution. from spitsbergen, w greenland, around iceland and along the coast of norway into the kattegat, including the øresund. also found at localities off ireland. occurrence. the arctic, subarctic, boreal and lusitanian regions. habitat. found to about 200 m on algae. subfossil finds. none. onoba semicostata (montagu 1803) distribution. around iceland, along the coast of norway and south to the mediterranean. absent from the eastern north sea coasts, but extends through the limfjord and kattegat, including the øresund, into the bælt sea and the most saline parts of the baltic (fretter & graham 1978b). occurrence. the subarctic, boreal and lusitanian regions. habitat. from the intertidal region to depths of 100 m. found under stones, amongst weeds, mussels and tunicates in shelly gravel, but only where there are quantities of silt (fretter & graham 1978b). subfossil finds. the bælt sea, kattegat, limfjord and vendsyssel regions, holocene. onoba proxima (forbes & hanley 1850) distribution. from the western coast of britain south to the mediterranean. occurrence. the lusitanian region. habitat. from 10 to 170 m on bottoms of muddy sand. only subfossil finds. the limfjord region, holocene. onoba vitrea (montagu 1803) fig. 10 distribution. from norway off lofoten and south to the mediterranean, extends through skagerrak into the kattegat, including the øresund, but absent from the limfjord. occurrence. the boreal and lusitanian regions. habitat. muddy bottoms at depths of 10–50 m in the northern parts of its range, but extending to 120 m in the south. further notes on onoba species in fretter & graham (1978b, p. 170). subfossil finds. the kattegat, limfjord, north sea, vendsyssel and skagen areas, holocene. in the skagen well recorded from the atlantic, the subboreal and the subatlantic. during the eemian recorded from the north sea. rissoa albella lovén 1846 fig. 11 distribution. from norway off the lofoten islands, and south to the mediterranean, extending into the limfjord and the kattegat with the danish fjords, including the øresund and the bælt sea. occurrence. the boreal and lusitanian regions. habitat. on rocky shores amongst weeds, sublittoral to 15 m. it is tolerant of some brackishness (fretter & graham 1978b). subfossil finds. the bælt sea, baltic, kattegat, limfjord, north sea, vendsyssel and skagen well from the subfig. 10. onoba vitrea (montagu 1803). skagen 3, 67.0–67.25 m b.s., lab. no. 720,93. × 9.6. mguh 25320. geus bulletin no 3.pmd 28-06-2004, 08:4533 34 atlantic, holocene. from the eemian recorded from the north sea. rissoa inconspicua alder 1844 distribution. from northern norway, north of lofoten, and south to the mediterranean. occurring in the limfjord, øresund and bælt sea. occurrence. the boreal and lusitanian regions. fretter & graham (1978b, p. 200) indicate that the species is found to the arctic. however, it is not recorded from iceland and the faroes but only from norway north of lofoten (thorson 1941, table ii, p. 141). habitat. typically sublittoral living on algae, and on sandy gravel to depths of about 100 m. tolerant of slightly brackish conditions. subfossil finds. the bælt sea, baltic, kattegat, limfjord, north sea and vendsyssel, holocene. during the eemian recorded from the bælt sea, kattegat and north sea regions. rissoa membranacea (j. adams 1800) distribution. from norway off lofoten and south to the canary islands. extends into the limfjord, kattegat, bælt sea and the westernmost part of the baltic (the rügen island). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone sublittorally to about 15 m associated with zostera or on weeds with the same habit, extending into brackish water. subfossil finds. the bælt sea, baltic, kattegat, limfjord, north sea and vendsyssel areas, holocene. during the eemian in the bælt sea and north sea regions. rissoa parva (da costa 1779) distribution. from norway, north of lofoten and the faeroes, south to the mediterranean, found in the limfjord (petersen 1986a) and the northern part of the øresund but not in the baltic according to fretter & graham (1978b). however, bondesen (1975) includes the species in the baltic, but excludes it from the bælt sea. petersen (1888, p. 93) regarded the species as being limited to the central part of the kattegat, depending on sufficient salt content – therefore the information of the occurrences in the baltic given by bondesen (1975) is surprising. occurrence. the boreal and lusitanian regions. habitat. from the tidal zone to about 25 m on fronds, smaller weeds and under stones. in the faeroes from the rock pools and the beach to a depth of 20 m (spärck & thorson 1931). subfossil finds. the limfjord and vendsyssel areas, holocene. during the eemian recorded from the bælt sea, north sea and vendsyssel regions. rissoa violacea desmarest 1814 fig. 12 distribution. rissoa violacea is not at all recorded from iceland (thorson 1941) and the faroes (spärck & thorson 1931) but from norway off the lofoten islands and south. both of the recent subspecies are recorded from the kattegat, including the øresund region. bonfig. 12. rissoa violacea desmarest 1814. skagen 4, 27.0– 27.5 m b.s., lab. no. 352,93. × 20. mguh 25322. fig. 11. rissoa albella lovén 1846. skagen 4, 30.0–30.5 m b.s., lab. no. 355,93. × 20. mguh 25321. geus bulletin no 3.pmd 28-06-2004, 08:4534 35 desen (1975) has rissoa violacea sensu lato from the skagerrak, the kattegat and the limfjord regions. occurrence. the boreal and lusitanian regions. habitat. in the tidal zone to about 50 m on weeds and amongst sandy gravel. subfossil finds. the limfjord, north sea, vendsyssel and skagen areas, holocene. in skagen recorded from the subatlantic. the species occurred during the eemian in the kattegat region. assiminea grayana fleming 1828 distribution. the species is confined to the north sea coasts. indanishwaters it extendssouth toblåvandshuk. occurrence. the boreal region. habitat. the species is limited to the upper parts of the salt-marsh areas on the vegetation. subfossil finds. none. caecum glabrum (montagu 1803) distribution. from norway off lofoten south to the mediterranean. it extends into kattegat, where it has been reported from the northern part (jensen & knudsen 1995), but not from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. sublittorally to about 250 m on sandy and sandy–muddy bottoms. subfossil finds. the limfjord, north sea and vendsyssel regions, holocene. recorded from the bælt sea and the north sea during the eemian. tornus exquisitus (jeffreys 1883) distribution. within the danish waters recorded from northern kattegat (jensen & knudsen 1995), but no record is given by fretter & graham (1978b, p. 232) so the species is not treated further. habitat. unknown for this species. subfossil finds. none. bittium reticulatum (da costa 1778) fig. 13 distribution. from off lofoten in norway south to the mediterranean. it is found through the skagerrak, limfjord, kattegat, including the øresund, and into the bælt sea, but not from the southern coastal part of the north sea. occurrence. the boreal and lusitanian regions. habitat. common in shallow sublittoral water, but recorded to 250 m. found on soft bottoms in association with weeds. subfossil finds. the species is recorded from all regions during the holocene. from skagen recorded from the subatlantic. from the eemian found in the bælt sea, kattegat, north sea, and vendsyssel regions. turritella communis risso 1826 fig. 14 distribution. from the lofoten islands south to the mediterranean. the species extends into all the danish waters as far as the øresund. occurrence. the boreal and lusitanian regions. habitat. from 10 to 200 m depths on soft bottoms. subfossil finds. the limfjord, north sea, vendsyssel and skagen areas, holocene. in the skagen area refig. 13. bittium reticulatum (da costa 1778). geus collection. vejlager xiv, denmark. × 9.6. mguh 25323. geus bulletin no 3.pmd 28-06-2004, 08:4535 36 corded from the subboreal and subatlantic. during the eemian recorded from the baltic, kattegat, north sea and vendsyssel regions. turritella erosa couthouy 1838 distribution. west and east greenland (thorson 1944b, p. 40), svalbard and on the northern coast of russia (macpherson 1971). occurrence. the arctic and subarctic regions. habitat. from 10 to 350 m on soft bottoms. subfossil finds. recorded from the vendsyssel region in early/middle weichselian. aporrhais pespelicani (linnaeus 1758) fig. 15 distribution. on the southern and western part of iceland, norway off lofoten with a questionable occurrence north of lofoten (thorson 1941), and south to the mediterranean. it does not occur off the west coast of denmark nor in its fjords, except the limfjord. it extends through the kattegat, including øresund (fretter & graham 1981, p. 297). empty shells are found in kieler bucht (arntz et al. 1976). occurrence. the boreal and lusitanian regions. habitat. sublittoral to depths of 180 m on mud, muddy sand and sand. subfossil finds. the baltic, limfjord, north sea, vendsyssel and skagen areas, holocene. in the skagen region from the subatlantic. during the eemian recorded from the bælt sea, baltic, north sea, and vendsyssel regions. aporrhais serresianus (michaud 1828) distribution. from southern and western iceland, norway off lofoten and south to the mediterranean. according to fretter & graham (1981) absent from the skagerrakand all danish seas; however, jensen & knudsen (1995) note a single record from the central kattegat. occurrence. the boreal and lusitanian regions. habitat. from sublittoral (as aporrhais pespelicani) down to 1000 m on finer muds. subfossil finds. none. crepidula fornicata (linnaeus 1758) distribution. c. fornicata is a late immigrant to europe (first recorded in the british isles late in the 19th century). the present distribution is from norway south to portugal. it reached the limfjord in 1934. in 1949– 50 it was found in the northern part of kattegat but has not spread further south (jensen & knudsen 1995). fig. 15. aporrhais pespelicani (linnaeus 1758). skagen 3, 0.0–30.0 m b.s. (washed sample). × 4.8. mguh 25325. fig. 14. turritella communis risso 1826. skagen 3, 73.10–73.30 m b.s., lab. no. 722,93. × 4.8. mguh 25324. geus bulletin no 3.pmd 28-06-2004, 08:4536 37 occurrence. the boreal–lusitanian regions within its european distribution. crepidula fornicata was transferred by man as seen also for species like mya arenaria. habitat. sublittoral to depths of c. 10 m. the animals live in chains, the oldest attached to a substrate which might be an oyster. the species was actually transported to europe with oysters (fretter & graham 1981, p. 311). subfossil finds. obviously none. capulus ungaricus (linnaeus 1758) distribution. empty shells recorded from sw and nw iceland and living specimens from norway north of lofoten south to the mediterranean. the record from greenland mentioned in fretter & graham (1981) cannot be sustained in the literature (thorson 1944b, 1951). occurrence. the boreal and lusitanian regions. habitat. usually sublittorally to 805 m attached to stones or the host animal. subfossil finds. none (young date from the north sea). lamellaria perspicua (linnaeus 1758) distribution. sw and nw iceland, norway from lofoten and south to the mediterranean. it occurs in the skagerrak but not on the danish coasts. occurrence. the boreal and lusitanian regions. habitat. from the tidal zone and downwards to depths of 1200 m, especially in the southern parts of its range on rocky shores and under stones. subfossil finds. none. velutina plicatilis (müller 1776) distribution. from east greenland and spitsbergen, southern and western iceland, norway north of lofoten and south to northern spain. it extends into the skagerrak, kattegat and the northern part of the north sea. occurrence. the arctic, subarctic, boreal and lusitanian regions. habitat. from 10 to c. 375 m deep on hard bottoms, usually in association with ascidians and hydroids. subfossil finds. none. velutina velutina (müller 1776) distribution. from spitsbergen, e and w greenland, around iceland and norway south to the mediterranean. it extends into the skagerrak and kattegat, including øresund. occurrence. arctic, subarctic, boreal and lusitanian regions. habitat. sublittoral extending to 1000 m on hard bottoms associated with tunicates. subfossil finds. none. trivia arctica (pulteney 1799) distribution. from norway off lofoten and south to the mediterranean. it extends into the skagerrak and kattegat (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. sublittoral to about 100 m, in southerly latitudes to about 1000 m. it is associated with ascidians. subfossil finds. none. trivia monacha (da costa 1778) distribution. from the british isles south to the mediterranean. the recent distribution in the north sea is questioned, and the species is not recorded from scandinavia (fretter & graham 1981, p. 329). occurrence. the lusitanian region. habitat. on rocky shores and under stones associated with ascidians. subfossil finds. the vendsyssel region, holocene. geus bulletin no 3.pmd 28-06-2004, 08:4537 38 amauropsis islandicus (gmelin 1791) distribution. from spitsbergen, e and w greenland, around iceland and norway. it extends into the skagerrak and kattegat, including øresund, although it is rare there (jensen & knudsen 1995). occurrence. the arctic, subarctic and boreal regions. habitat. sublittorally to about 80 m deep on sandy clay bottoms. subfossil finds. none (late date from the north sea). lunatia alderi (forbes 1838) fig. 16 distribution. from southern and western iceland and norway off the lofoten islands south to the mediterranean (thorson 1941). it extends into the skagerrak, kattegat and øresund, but the species is not recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. sublittoral between 10 and 50 m, extending to 2000 m. infaunal on sandy shores, clean sand and some admixture of mud. according to petersen (1888), it is common on mixed bottoms in the kattegat. subfossil finds.the limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen area recorded from the atlantic, subboreal and subatlantic. from the eemian recorded from the baltic, kattegat, north sea, vendsyssel and skagen regions. lunatia catena (da costa 1778) distribution. from the skagerrak and kattegat, including øresund but not the limfjord, south to the mediterranean (fretter & graham 1981, p. 339). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone down to about 125 m on sandy bottoms. subfossil finds. the limfjord, north sea and vendsyssel regions, holocene. lunatia montagui (forbes 1838) fig. 17 distribution. from w and s iceland, norway north of the lofoten islands, and south to the mediterranean. it occurs in the skagerrak and kattegat, including øresund, however, rare in the sound as the presiding species according to jensen & knudsen (1995). occurrence. the boreal and lusitanian regions. habitat. from 15 to 200 m depth on sandy and muddy bottoms. subfossil finds. recorded from the skagen well from the subatlantic. fig. 16. lunatia alderi (forbes 1838). skagen 4, 26.0–26.5 m b.s., lab. no. 351,93. × 20. mguh 25326. fig. 17. lunatia montagui (forbes 1838). skagen 4, 20.0–20.5 m b.s., lab. no. 345,93. × 4.8. mguh 25327. geus bulletin no 3.pmd 28-06-2004, 08:4538 39 lunatia pallida (broderip & sowerby 1829) distribution. from spitsbergen, e and w greenland, around iceland and norway, south to the north sea, skagerrak and the kattegat, including the øresund. occurrence. the arctic, subarctic and boreal regions. habitat. from 10 to 2000 m on clay bottoms – the greatest depths in the most southerly parts of its range (fretter & graham 1981). in greenland waters commonin the arctic macoma community (thorson 1944b). subfossil finds. the vendsyssel region during the early/ middle weichselian and late weichselian (the older and younger yoldia sea respectively). natica affinis (gmelin 1790) distribution. from spitsbergen, w and e greenland, around iceland and norway south to the mediterranean. however, within the lusitanian region the species lives in deep water (fretter & graham 1981, p. 345). the species is recorded from danish waters (jensen & knudsen 1995). occurrence. the arctic, subarctic, boreal and lusitanian regions. habitat. from about 4 m depth in high latitudes to well over 2000 m in low ones on sandy, muddy and clay bottoms. as mentioned for other species with a wide geographical distribution this is a ‘tropical submerge’ which is quite common for cold-water animals which in the northern regions inhabit the surface water to occur mainly or exclusively in deeper zones in the southern seas (ekman 1953, p. 112). the species is found in the arctic macoma community (spärck 1937). subfossil finds. recorded from the vendsyssel area during the early/middle weichselian (the older yoldia sea), and the late weichselian (the younger yoldia sea). order heterogastropoda triphora adversa (montagu 1803) distribution. from norway off the lofoten islands and south to spain. it extends into the kattegat, including øresund and the bælt sea, but not recorded from the limfjord. occurrence. the boreal and lusitanian regions. habitat. sublittorally to 100 m under stones, algae or associated with sponges. subfossil finds. the bælt sea, kattegat, limfjord, north sea and vendsyssel areas, holocene. recorded from the bælt sea and the north sea during the eemian. cerithiella metula (lovén 1846) distribution. s and w iceland, norway north of lofoten, and south to the mediterranean. recorded from the danish waters (jensen & knudsen 1995), although rare in the skagerrak and not occurring in the kattegat (fretter & graham 1982, p. 377). occurrence. the boreal and lusitanian regions. habitat. from 40 to 400 m depth on soft bottoms. subfossil finds. none. cerithiopsis barleei jeffreys 1867 distribution. according to fretter & graham (1982), from sw england south to the mediterranean. however, the species is recorded from danish waters, i.e. the øresund area, although as rare (jensen & knudsen 1995), but is not found in the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. sublittorally associated with sponges. subfossil finds. the limfjord region, holocene. geus bulletin no 3.pmd 28-06-2004, 08:4539 40 cerithiopsis tubercularis (montagu 1803) distribution. from norway off the lofoten islands and south to the mediterranean. recorded from the swedish west coast, but not in danish waters (fretter & graham 1982). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone and sublittorally to 100 m. the species is found on sponges. only subfossil finds. the limfjord region, holocene, and the north sea during the eemian. epitonium clathratulum (kanmacher 1797) distribution. from norway and south to the mediterranean. it extends into the kattegat, including the øresund. occurrence. the boreal and lusitanian regions. habitat. from 30 to 100 m deep on sandy–muddy bottoms. subfossil finds. none. epitonium clathrus (linnaeus 1758) distribution. from norway off the lofoten islands and south to the mediterranean. the species extends into the kattegat and øresund (jensen & knudsen 1995), but does not enter the danish fjords (fretter & graham 1982, p. 387). occurrence. the boreal and lusitanian regions. habitat. sublittorally from 5 to 70 m on sandy–muddy bottoms. subfossil finds. the limfjord, north sea and vendsyssel regions, holocene. recorded from the bælt sea and the north sea during the eemian. epitonium trevelyanum (johnston 1841) fig. 18 distribution. from norway off the lofoten islands and south to the mediterranean. the species extends into the kattegat, including the øresund. occurrence. the boreal and lusitanian regions. habitat. from 30 to 200 m depth on sandy–muddy bottoms. subfossil finds. the skagen area, holocene, recorded from the subboreal and the subatlantic. epitonium turtonis (turton 1819) distribution. from norway off the lofoten islands (thorson 1941) and south to the mediterranean. occurring in the kattegat (fretter & graham 1982) and øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from 5 to 20 m deep on sandy–muddy bottoms. subfossil finds. the limfjord and vendsyssel regions, holocene. aclis ascaris (turton 1819) distribution. from norway off lofoten and south to the mediterranean. the species extends into the skagerrak and kattegat, including the øresund. occurrence. the boreal and lusitanian regions. habitat. from 10 to 50 m deep on soft sandy bottoms. subfossil finds. the north sea, holocene. fig. 18. epitonium trevelyanum (johnston 1841). skagen 3, 58.5– 60.0 m b.s., lab. no. 98,93. × 9.6. mguh 25328. geus bulletin no 3.pmd 28-06-2004, 08:4540 41 aclis minor (brown 1827) fig. 19 distribution. from norway off the lofoten islands and south to the mediterranean. the species is recorded from the southern kattegat and øresund. occurrence. the boreal and lusitanian regions. habitat. from 15 to 150 m deep on bottoms of sand, muddy sand or gravel. subfossil finds. the limfjord, north sea and skagen areas, holocene; in the skagen well recorded from the atlantic, subboreal and subatlantic. aclis walleri jeffreys 1867 distribution. from norway off the lofoten islands and south to the mediterranean. recorded from north of skagen (petersen 1888) and the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. according to fretter & graham (1982), at greater depths than the other aclis ssp. – down to 550 m on soft bottoms. subfossil finds. the north sea, holocene. eulima bilineata (alder 1848) distribution. from norway north of lofoten, south to the mediterranean. in the danish waters recorded from the southern kattegat. occurrence. the boreal and lusitanian regions. habitat. from 20 to 250 m deep on soft bottoms associated with ophiuroids. subfossil finds. none. haliella stenostoma (jeffreys 1858) distribution. off west greenland, around iceland, norway off the lofoten islands, and south to the mediterranean. occurrence. the subarctic, boreal and lusitanian regions. habitat. sublittoral from about 70 to 3000 m on soft bottoms. the species has its main occurrence at rather great depths (thorson 1941), but is also recorded from danish waters (jensen & knudsen 1995). subfossil finds. none. polygireulima sinuosa (sacco 1836) fig. 20 distribution. from the kattegat and south to the mediterranean. occurrence. the boreal and lusitanian regions. habitat. sublittoral from 30 to 150 m on soft bottoms. subfossil finds. the skagen area, holocene, recorded from the subatlantic. fig. 19. aclis minor (brown 1827). skagen 3, 55.1–55.3 m b.s., lab. no. 716,93. × 20. mguh 25329. fig. 20. polygireulima sinuosa (sacco 1836). skagen 3, 38.19–38.24 m b.s., core sample k-6. × 9.6. mguh 25330. geus bulletin no 3.pmd 28-06-2004, 08:4541 42 polygireulima monterosatoi (monterosato 1890) distribution. norway south of the lofoten islands, and south to the mediterranean. occurrence. the boreal and lusitanian regions. habitat. from 20 to 120 m deep on sandy muddy or gravelly bottoms. however, as fretter & graham (1982, p. 421) say: “presumably these animals attack echinoderms like their relatives, but which is not known”. subfossil finds. none. vitreolina collensi (sykes 1903) fig. 21 distribution. from the west coast of britain and ireland to the mediterranean. fig. 21. vitreolina collensi (sykes 1903). skagen 3, 70.10–70.30 m b.s., lab. no. 721,93. × 20. mguh 25331. fig. 22. a, b: vitreolina philippii (rayneval & pouzi 1854). skagen 4, 21.0–21.5 m b.s., lab. no. 346,93. × 20. mguh 25332. occurrence. the lusitanian region. habitat. sublittorally to 35–40 m on soft bottoms. only subfossil finds. the skagen area, holocene, recorded from the subboreal and the subatlantic. vitreolina philippii (reyneval & ponzi 1854) fig. 22a, b distribution. from norway off the lofoten islands and south to the mediterranean. according to petersen (1888), very common in danish waters from the skagerrak, the kattegat and the øresund, and the bælt sea, but not recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from the lowest part of the tidal zone to a depthof 200 m on soft bottoms. fretter & graham (1982, p. 422) hold it as perhaps the most common local eulimid and like other eulimids an intermittent parasite of echinoderms. subfossil finds. the limfjord, vendsyssel and skagen areas, holocene, recorded from the subatlantic in the skagen well. occurring in the eemian in the vendsyssel region. graphis albida (kanmacher 1798) fig. 23 distribution. from southern norway south to the mediterranean. the species has a record from the limfjord (petersen 1986a). geus bulletin no 3.pmd 28-06-2004, 08:4542 43 occurrence. the boreal and lusitanian regions. habitat. from low in the tidal zone to 30 m deep on muddy and sandy bottoms. only subfossil finds. the skagen well area, holocene, recorded from the subatlantic. melanella lubrica (monterosato 1891) fig. 24 distribution. from norway south to iberia. the species is recorded from the danish waters south into the øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from 14 to 100 m on soft bottoms of muddy sand and gravel. the species is an intermittent ectoparasite (fretter & graham 1982). subfossil finds. the skagen area, holocene, recorded from the atlantic and subatlantic. melanella alba (da costa 1778) fig. 25 distribution. from norway off the lofoten islands and south to the mediterranean. occurrence. the boreal and lusitanian regions. habitat. from 16 to 135 m deep on muddy sand and gravel bottoms, an ectoparasite of holothurians. only subfossil finds. the skagen well area, holocene, recorded from the subboreal. hemiaclis ventrosa (jeffreys ms fricle 1874) fig. 26 distribution. from west and south iceland, norway off the lofoten islands, and south to the bay of biscay. occurrence. the boreal and lusitanian regions. habitat. from 100 to 300 m deep on soft bottoms. only subfossil finds. the skagen well area, holocene, recorded from the subatlantic. fig. 23. graphis albida (kanmacher 1798). skagen 4, 25.0–25.5 m b.s., lab. no. 350,93. × 40. mguh 25333. fig. 24. melanella lubrica (monterosato 1891). skagen 3, 83.73– 83.83 m b.s., lab. no. 512,93. × 40. mguh 25334. fig. 25. melanella alba (da costa 1778). skagen 3, 74.89–75.00 m b.s., lab. no. 509,93. × 40. mguh 25335. geus bulletin no 3.pmd 28-06-2004, 08:4543 44 pelseneeria stylifera (turton 1826) distribution. from norway off the lofoten islands and south to the mediterranean. few recorded from the skagerrak and kattegat, including øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. the animals are confined to the surface of regular sea urchins (fretter & graham 1982, p. 431). subfossil finds. none. enteroxenos oestergreni bonnevie 1902 distribution. recorded from scandinavia, a single danish record (jensen & knudsen 1995). occurrence. the boreal region. habitat. a parasite in the holothurian stichopus tremulus (jensen & knudsen 1995). subfossil finds. none. order neogastropoda nucella lapillus (linnaeus 1758) distribution. w greenland, around iceland, norway from north of lofoten, and south to the straits of gibraltar. the species reaches into the skagerrak, but it is uncommon in danish waters (fretter & graham 1984). however, it occurs on breakwaters along the north sea and skagerrak coasts. there are a few records from the southern kattegat and øresund (jensen & knudsen 1995). occurrence. the subarctic, boreal and lusitanian regions. habitat. intertidal on rocky shores and extends, albeit rarely, to depths of 30 to 40 m. it avoids very weedy shores and seems to stand only limited reduction of salinity (fretter & graham 1984, p. 445). subfossil finds. the kattegat, limfjord (exposed towards the skagerrak), north sea and vendsyssel regions, holocene. boreotrophon clathratus (linnaeus 1767) distribution. spitsbergen, e and w greenland, around iceland, the faroes and the coast of norway south to the skagerrak and kattegat. occurrence. the arctic, subarctic and boreal regions. habitat. from 8 m to over 1000 m on soft bottoms. subfossil finds. the vendsyssel area, late weichselian (the younger yoldia sea). boreotrophon truncatus (ström 1768) distribution. spitsbergen, e and w greenland, around iceland, norway south to the biscay. the species extends in to danish waters south to the øresund. occurrence. the arctic, subarctic, boreal and lusitanian regions. habitat. from the laminarian zone to depths of about 200 m on bottoms of a stony, gravelly or muddy nature. subfossil finds. none. ocenebra erinacea (linnaeus 1758) distribution. from the southern coasts of britain south to the mediterranean. however, the few records from fig. 26. hemiaclis ventrosa (jeffreys ms fricle 1874). skagen 4, 30.0–30.5 m b.s., lab. no. 355,93. × 20. mguh 25336. geus bulletin no 3.pmd 28-06-2004, 08:4544 45 danish waters might have been introduced with oysters (jensen & knudsen 1995). occurrence. the lusitanian region. habitat. sublittoral to 150 m deep on stony bottoms. subfossil finds. none. trophonopsis barvicensis (johnston 1825) distribution. w and s iceland, norway and south to the british isles, and further south (france) at greater depths. in danish waters from the kattegat, including the øresund, although rare (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. sublittoral at a few metres’ depth at the northern end of its range to 300–400 m at the southern end. subfossil finds. none. buccinum undatum linnaeus 1758 fig. 27 distribution. spitsbergen, w greenland, around iceland, norway and south to the bay of biscay (thorson 1944b). the species extends into the kattegat, limfjord, and the bælt sea with the mecklenburger bucht as the easternmost position. occurrence. the subarctic, boreal and lusitanian regions. habitat. sublittorally to about 1200 m deep usually on soft bottoms. subfossil finds. the kattegat, limfjord, north sea, vendsyssel and skagen areas, holocene, recorded as a fragment from the skagen well from the subatlantic. during the eemian in the kattegat, north sea, and vendsyssel regions. from the vendsyssel area found during the late weichselian (the younger yoldia sea). buccinum cyaneum bruguière 1792 distribution. spitsbergen, w and e greenland around iceland and norway north of lofoten. occurrence. the arctic, subarctic and northern part of the boreal regions. habitat. from 0 to 392 m on all sorts of bottoms – sand, clay, stones and algae (thorson 1944b). only subfossil finds. from the vendsyssel area recorded during the late weichselian (the younger yoldia sea). colus gracilis (da costa 1778) distribution. s and w iceland (empty shells), norway off the lofoten islands and south to portugal. the species extends into the kattegat. occurrence. the boreal and lusitanian regions. habitat. usually from 30 to 800 m deep (less common and deeper in the south). subfossil finds. none. colus jeffreysianus (fischer 1868) distribution. from norway south to the mediterranean. the species extends through the skagerrak to the kattegat, including the northern part of the øresund. occurrence. the boreal and lusitanian regions. habitat. from 30 to 2000 m deep on soft bottoms. subfossil finds. none.fig. 27. buccinum undatum linnaeus 1758. geus collection. limfjord, denmark. height 70 mm. mguh 25337. geus bulletin no 3.pmd 28-06-2004, 08:4545 46 colus sabini (gray 1824) distribution. from w greenland (empty shells) and s and w iceland, northern north sea and extending into the skagerrak (jensen & knudsen 1995). occurrence. the boreal region. (a subarctic extension is not considered, since only empty shells have been found off w greenland and no occurrences on n and e iceland (thorson 1941, 1944b).) habitat. from 35 to 1500 m deep on muddy bottoms. subfossil finds. none. liomesus ovum (turton 1825) distribution. from greenland, the faeroes and the coasts of norway. according to fretter & graham (1984, p. 465), the species is not recorded from the skagerrak or kattegat; however, jensen & knudsen (1995) mentioned this species as occurring in danish waters. furthermore, the species liomesus ovum cannot be found in thorson (1944b), who has buccinum ovum middendorff, which is not the same according to fretter & graham (1962), and the occurrence off the faeroes cannot be confirmed in spärck & thorson (1931). occurrence. the boreal region with only uncertain outposts into the subarctic and lusitanian regions. habitat. from 70 to 400 m deep on soft bottoms. subfossil finds. none. neptunea antiqua (linnaeus 1758) distribution. from southern norway south to the bay of biscay. the species extends into the kattegat, øresund (jensen & knudsen 1995), and the bælt sea as far east as lübecker bucht. the species has also been recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from 15 to 1200 m on all kinds of bottom, mainly soft. subfossil finds. the vendsyssel area, holocene. neptunea despecta (linnaeus 1758) distribution. from spitsbergen, e and w greenland, around iceland and the coasts of norway, south to the seas off denmark according to fretter & graham (1984), but not mentioned by jensen & knudsen (1995) in their annotated check list of recent marine molluscs of danish waters. occurrence. the arctic, subarctic and boreal regions. habitat. from 6 to 1400 m on soft bottoms. only subfossil finds. the vendsyssel region during the late weichselian (the younger yoldia sea). turrisipho moebii (dunker & metzger 1874) distribution. from the coasts of northern and southern norway and the faeroes (sipho sarsi in spärck & thorson 1931). the species extends into the skagerrak. occurrence. the boreal region. habitat. from 200 m to 1000 m deep on soft bottoms. subfossil finds. none. hinia incrassata (ström 1768) distribution. s and w iceland, norway north of lofoten, and south to the mediterranean (thorson 1941). the species extends into the kattegat and øresund, although rare in this place (jensen & knudsen 1995), and it has not been found in the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. rocky coasts in the lower part of the tidal zone. mainly found in the shallow sublittoral, but may extend to about 200 m. subfossil finds. the limfjord and vendsyssel regions, holocene. recorded from the vendsyssel region during the eemian. hinia pygmaea (lamarck 1822) fig. 28 distribution. from norway off the lofoten islands and geus bulletin no 3.pmd 28-06-2004, 08:4546 47 south to the mediterranean. it extends into danish waters such as the skagerrak, limfjord and kattegat and øresund. occurrence. the boreal and lusitanian regions. habitat. from 1 m to about 200 m on sandy bottoms. subfossil finds. the limfjord, north sea, vendsyssel and skagen areas, holocene. recorded from the subatlantic in the skagen well material. found in the bælt sea, kattegat and north sea regions during the eemian. hinia reticulata (linnaeus 1758) fig. 29 distribution. from norway off lofoten and south to the mediterranean. in danish waters within the limfjord and kattegat with fjords, the øresund and the bælt sea. occurrence. the boreal and lusitanian regions. habitat. from the tidal zone to about 15 m deep on soft bottoms. subfossil finds. the bælt sea, baltic, kattegat, limfjord, north sea, vendsyssel and skagen regions, holocene. in the skagen well material recorded from the subatlantic. from the bælt sea, baltic, kattegat, north sea and vendsyssel regions recorded from the eemian. troschelia bernicensis (king 1846) distribution. from norway north of lofoten and south to the west coast of scotland and the dogger bank. recorded fromdanishwaters by jensen&knudsen (1995). occurrence. the boreal and lusitanian (northern part) regions. habitat. lives on the continental shelves and upper slopes, between 90 and 2700 m (poppe & goto 1991). subfossil finds. none. cytharella coarctata (forbes 1840) distribution. from norway off the lofoten islands (thorson 1941: mangelia costata) and south to the mediterranean. the species is recorded from the skagerrak and kattegat where it extends into the øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone to 250 m deep on sandy bottoms. subfossil finds. the vendsyssel area, holocene. during the eemian in the north sea region. oenopota incisula (verrill 1882) distribution. the species is known from the boreal zone of the east coast of north america, and is comfig. 28. hinia pygmaea (lamarck 1822). skagen 4, 12.0–12.5 m b.s., lab. no. 337,93. × 9.6. mguh 25338. fig. 29. hinia reticulata (linnaeus 1758). skagen 3, 55.10–55.30 m b.s., lab. no. 716,93. × 4.8. mguh 25339. geus bulletin no 3.pmd 28-06-2004, 08:4547 48 mon off west greenland (posselt & jensen 1898), but it is not recorded from iceland or the coasts of the north-east atlantic (thorson 1941). occurrence. the arctic, subarctic, and in north america into the boreal regions. habitat. in the canadian north-east region the species has been collected from 6–7 to 140 m deep on clay (macpherson 1971). only subfossil finds. from the vendsyssel region recorded during the eemian and early/middle weichselian (the older yoldia clay). oenopota trevelliana (turton 1834) distribution. spitsbergen, e and w greenland, around iceland (thorson 1941: bela trevelliana (turton)), norway from north of lofoten, and south to the british isles. the species extends south to the kattegat, including øresund. occurrence. the arctic, subarctic, boreal and lusitanian (northern part) regions. habitat. sublittorally from 25 m to depths over 300 m on fine sand. subfossil finds. the vendsyssel area during the eemian. oenopota turricola (montagu 1803) fig. 30 distribution. greenland, iceland, faeroes, norway south to scotland (spärck & thorson 1931). the species extends into the limfjord and kattegat (fretter & graham 1984) and the øresund (jensen & knudsen 1995). occurrence. the arctic, subarctic and boreal regions. habitat. sublittoral from 20 to 200 m on sandy bottoms. subfossil finds. the limfjord, vendsyssel and skagen areas, holocene. from the skagen well recorded from the subatlantic. found in the vendsyssel area from the late weichselian (the younger yoldia sea). oenopota violacea (mighels & adams 1842) distribution. e and w greenland, spitsbergen, around iceland and south along the coast of norway, but not reaching the british isles (thorson 1941). occurrence. the arctic, subarctic and boreal regions. habitat. from 1 m to 761 m on mud and stones. only subfossil finds. the vendsyssel area during the eemian. bela exarata g.o. sars 1878 distribution. e and w greenland, spitsbergen, around iceland, the faeroes, and norway from north of lofoten and west of ireland. the species is also recorded west of iceland, where it has been found at depths down to 2214 m (thorson 1941). occurrence. the arctic, subarctic, boreal regions with lusitanian outposts. habitat. in the northern part of its range it belongs to the shallow-water species (norway and e greenland from 3 m) (thorson 1941). only subfossil finds. from the eemian in the vendsyssel area. mangelia attenuata (montagu 1803) distribution. from norway off the lofoten islands and fig. 30. oenopota turricola (montagu 1803). skagen 4, 21.0– 21.5 m b.s., lab. no. 346,93. × 9.6. mguh 25340. geus bulletin no 3.pmd 28-06-2004, 08:4548 49 south to the mediterranean. the species extends into the skagerrak and kattegat, including the øresund. occurrence. the boreal and lusitanian regions. habitat. from 5 to 150 m deep on sand or clay bottoms. subfossil finds. none. mangelia brachystoma (philippi 1844) fig. 31 distribution. from norway off the lofoten islands and south to the mediterranean. the species extends into the skagerrak and kattegat, including the øresund. occurrence. the boreal and lusitanian regions. habitat. sublittorally from 4 m to 60 m deep on bottoms of sand and sandy mud. subfossil finds. the skagen area, holocene, recorded from the subatlantic in the skagen well cores. during the eemian found in the vendsyssel region. mangelia nebula (montagu 1803) distribution. norway off the lofoten islands and south to the mediterranean. the species is recorded a few times from the kattegat and extends into the øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from 10 to 50 m deep on sandy bottoms. subfossil finds. none. raphitoma purpurea (montagu 1803) distribution. from northern norway off the lofoten islands south into the mediterranean. occurrence. the boreal and lusitanian regions. habitat. from 10 to 100 m deep on sandy, gravelly and stony bottoms. only subfossil finds. the limfjord region, holocene. raphitoma asperrima (brown 1827) distribution. from the coast of norway south to the mediterranean, extending into the kattegat, including the øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from 20 to 100 m deep on sandy bottoms. subfossil finds. none. raphitoma leufroyi (michaud 1821) distribution. from the coast of norway off the lofoten islands and south to the mediterranean, extending into the skagerrak and with a few records from the northern kattegat (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone to 150 m deep on sandy, shelly and stony bottoms. subfossil finds. none. raphitoma linearis (montagu 1803) distribution. sw and nw iceland, norway from north of lofoten and south to the mediterranean. the species extends into the skagerrak and kattegat and øresund, although rare (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. fig. 31. mangelia brachystoma (philippi 1844). skagen 3, 58.04– 58.25 m b.s., lab. no. 717,93. × 9.6. mguh 25341. geus bulletin no 3.pmd 28-06-2004, 08:4549 50 habitat. from 10 to 200 m deep on sandy, shelly and stony bottoms. subfossil finds. the limfjord and vendsyssel area, holocene. during the eemian found in the vendsyssel area (the turritella terebra zone in the skærumhede sequence). taranis borealis bouchet & warén 1980 distribution. so far as known, this species is confined to waters off western norway and the skagerrak (fretter & graham 1984, p. 548). occurrence. the boreal region. habitat. from 150 m to nearly 2000 m deep on soft bottoms. subfossil finds. none. taranis moerchi (malm 1861) distribution. from norway north of lofoten south to the mediterranean, extending into the kattegat. occurrence. the boreal and lusitanian regions. habitat. from 80 m deep near the northern limits of its range to over 2000 m elsewhere on soft bottoms. subfossil finds. none. admete viridula (fabricius 1780) distribution. spitsbergen, e and w greenland, around iceland, empty shells from off the faeroes, norway from north of the lofoten islands, and south to the northern borders of the north sea. included in the recent danish fauna (jensen & knudsen 1995). occurrence. the arctic, subarctic and boreal regions. habitat. from a few metres to depths of 1000 m, the greatest depths in the south of its range (fretter & graham 1984, p. 507) on soft bottoms. subfossil finds. recorded from the early/middle weichselian (the portlandia arctica zone in the skærumhede sequence) in the vendsyssel region. subclass heterobranchia order heterostropha omalogyra atomus (philippi 1841) distribution. w greenland, around iceland and norway, south to the mediterranean. known from only a few places in danish waters. occurrence. the subarctic, boreal and lusitanian regions. habitat. from the lower part of the shore to a depth of 20 m, occurring on seaweeds. only subfossil finds. the limfjord region, holocene. brachystomia carozzai van aartsen 1987 distribution. from the southern part of norway (spärck & thorson 1931) and south to the mediterranean. the species extends into the kattegat and limfjord. as commented on by fretter et al. (1986, p. 605) and jensen & knudsen (1995), the determination of these small snails living ectoparasitically on other marine organisms is still in progress, so the actual situation for the record of subfossil material should be taken with great precaution. occurrence. the boreal and lusitanian regions. habitat. from the tidal zone, where it occurs in crevices, to depths of about 70 m. subfossil finds. none. brachystomia eulimoides hanley 1844 distribution. from norway off the lofoten islands and south to the mediterranean. recorded from the limfjord, according to jensen & knudsen (1995) the only danish record, following fretter et al. (1986, p. 602). occurrence. the boreal and lusitanian regions. habitat. most frequently found on living animals of pecten, chlamys, oysters and turritella to depths of 120 m. subfossil finds. the kattegat, limfjord, north sea and geus bulletin no 3.pmd 28-06-2004, 08:4550 51 vendsyssel regions, holocene. from the north sea region during the eemian. odostomia scalaris macgillivray 1843 distribution. from southern norway south to the mediterranean. the species extends into the limfjord, through the kattegat, including the øresund, and the bælt sea into the kiel bay (fretter et al. 1986, p. 600). occurrence. the boreal and lusitanian regions. habitat. associated primarily with banks of mytilus edulis, but also recorded from other hosts (fretter et al. 1986, p. 600). subfossil finds. the limfjord and vendsyssel regions, holocene. from the bælt sea and north sea during the eemian. chrysallida decussata (montagu 1803) fig. 32 distribution. mainly southern distribution, with the shetlands as the northernmost post, but recorded from the øresund (jensen & knudsen 1995) and there are older records from east of scotland (fretter et al. 1986). occurrence. the boreal and lusitanian regions. habitat. from 14 to 40 m deep on sandy and shelly bottoms. subfossil finds.thelimfjordandskagenareas,holocene, in the skagen well recorded from the subatlantic. chrysallida eximia (jeffreys 1849) distribution. sw and nw iceland, norway north of lofoten, and south to western scotland. there is no record from danish waters. occurrence. the boreal and lusitanian (northernmost) region. habitat. from 20 m to more than 1000 m, the greater depths in the southern part of its range, on soft gravelly bottoms. only subfossil finds. the limfjord region, holocene. from the vendsyssel area during the eemian. chrysallida indistincta (montagu 1808) distribution. from norway off the lofoten islands (thorson 1941) and south to the mediterranean. it extends into the kattegat and øresund (jensen & knudsen 1995), but is not recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from 7 to 100 m deep on sandy bottoms. subfossil finds. the limfjord, north sea and vendsyssel areas, holocene. recorded from the north sea during the eemian. chrysallida obtusa (brown 1827) distribution. from norway off the lofoten islands and south to the mediterranean. the species extends into the limfjord and through the kattegat into the øresund (fretter et al. 1986, p. 562). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone in rock pools to 90 m deep in stony places, associated with oysters (fretter et al. 1986, p. 562). subfossil finds. the limfjord and vendsyssel regions, holocene. from the bælt sea and north sea regions during the eemian. fig. 32. chrysallida decussata (montagu 1803). skagen 4, 15.0– 15.5 m b.s., lab. no. 340,93. × 20. mguh 25342. geus bulletin no 3.pmd 28-06-2004, 08:4551 52 chrysallida spiralis (montagu 1803) distribution. from norway north of lofoten and south to the mediterranean. according to fretter et al. (1986, p. 574), the species extends into the kattegat and øresund, but it is absent from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone to about 120 m deep, often abundant in the neighbourhood of tubes of sedentary polychaetes (fretter et al. 1986). subfossil finds. the kattegat, limfjord, north sea and vendsyssel regions, holocene. the bælt sea, kattegat, and north sea regions during the eemian. ebala nitidissima (montagu 1803) distribution. from south of norway to the mediterranean. recorded from the kattegat region with fjords and the øresund and the bælt sea regions as far as kiel bay (fretter et al. 1986, p. 630), but absent from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from 5 to 50 m deep on muddy sand or shelly bottoms. subfossil finds. the limfjord and vendsyssel regions, holocene. from the bælt sea and north sea areas during the eemian. eulimella laevis (brown 1827) distribution. from norway off the lofoten islands and south to the mediterranean. it has not been recorded from danish waters except the øresund region (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from 20 to 400 m deep on muddy sand. subfossil finds. the limfjord, north sea and vendsyssel regions, holocene. eulimella scillae (scacchi 1835) fig. 33 distribution. from norway off the lofoten islands, but empty shells only off sw iceland (thorson 1941), and south to the mediterranean. it extends into the kattegat along the swedish west coast (fretter et al. 1986), but occurs in the øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from 20 to 400 m deep on muddy sand or sand. subfossil finds. the limfjord and skagen areas, holocene, recorded in the skagen well from the atlantic and subboreal. from the vendsyssel area found during the eemian. ondina divisa (j. adams 1797) distribution. from s and w iceland, northern norway (w finmarken) (thorson 1941) and south to the biscay. the species extends through the kattegat to the øresund (fretter et al. 1986, p. 582). occurrence. the boreal and lusitanian regions. habitat. from 18 to 200 m deep on sandy and gravelly mud. subfossil finds. the limfjord region, holocene. from the vendsyssel area found during the eemian. fig. 33. eulimella scillae (scacchi 1835). skagen 3, 82.34–82.50 m b.s., lab. no. 729,93. × 9.6. mguh 25343. geus bulletin no 3.pmd 28-06-2004, 08:4552 53 ondina obliqua (alder 1884) distribution. according to fretter et al. (1986, p. 586) from southern scandinavia to biscay. however, spärck & thorson (1931) only mentioned scotland. furthermore, only western localities have been reported from the british isles (fretter et al. 1986), and it is doubtful whether the danish records actually refer to this species. therefore, the species is here considered to be lusitanian and should not be taken as present in the recent danish fauna, although recorded by jensen & knudsen (1995). habitat. from 30 to 60 m deep in gravelly or sandy mud. subfossil finds. none. ondina diaphana (jeffreys 1848) distribution. according to spärck & thorson (1931), present from southern norway and south to the mediterranean. in danish waters the kattegat and the øresund. occurrence. the boreal and lusitanian regions. habitat. from 20 to 90 m deep on soft bottoms. subfossil finds. the limfjord and north sea regions, holocene. liostomia clavula (lovén 1846) distribution. from southern norway south to the mediterranean. extends through the kattegat to the øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from 30 to 90 m deep on soft bottoms associated with pennatula (fretter et al. 1986, p. 590). subfossil finds. none. liostomia afzelii warén 1991 this species, newly established, will not be considered further. odostomia acuta jeffreys 1848 distribution. from norway north of lofoten and south to the mediterranean. the species is found in the skagerrak, kattegat and øresund (fretter et al. 1986, p. 613). occurrence. the boreal and lusitanian regions. habitat. from 20 to 30 m deep. perhaps associated with bryozoans. subfossil finds. the limfjord region, holocene. odostomia conoidea winckworth 1932 fig. 34 distribution. from norway off the lofoten islands and south to the mediterranean. recorded from the skagerrak (fretter et al. 1986) and the øresund (jensen & knudsen 1995), but not from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from 10 to 150 m deep, usually in association with the starfish astropecten irregularis (fretter et al. 1986, p. 617). subfossil finds. the bælt sea, limfjord, north sea, vendsyssel and skagen regions, holocene. in the skagen area recorded from the subatlantic. fig. 34. odostomia conoidea winckworth 1932. skagen 3, 48.90– 49.10 m b.s., lab. no. 714,93. × 40. mguh 25344. geus bulletin no 3.pmd 28-06-2004, 08:4553 54 odostomia turrita hanley 1844 distribution. from norway north of lofoten and south to the mediterranean. the species extends through the kattegat to the øresund (fretter et al. 1986, p. 612), but is not recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone to 100 m deep on weed and clay bottoms. subfossil finds. the limfjord and vendsyssel regions, holocene. recorded from the vendsyssel area during the eemian. odostomia albella lovén 1846 distribution. empty shells recorded from spitsbergen, sw and nw iceland, norway north of lofoten and south to the mediterranean. the species occurs in the skagerrak (fretter et al. 1986), the øresund area (jensen & knudsen 1995), and the limfjord region (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone to depths of about 100 m on boulders, associated with growth of pomatoceros (fretter et al. 1986). subfossil finds. the limfjord, north sea and vendsyssel regions, holocene. from the north sea region during the eemian. odostomia plicata (montagu 1803) distribution. from the southern part of scandinavia (not southern norway (spärck & thorson 1931)) and south to the mediterranean (fretter et al. 1986, p. 610). the species extends into the kattegat, including the øresund (jensen & knudsen 1995), but is not recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. in the tidal zone associated with pomatoceros triqueter (fretter et al. 1986). subfossil finds. the limfjord and vendsyssel regions, holocene. odostomia umbilicaris (malm 1863) fig. 35 distribution. from southern norway (spärck & thorson 1931) to the british isles. the species extends to the swedish west coast but not further into the kattegat (fretter et al. 1986, p. 620). occurrence. the boreal and lusitanian (northern part) regions. habitat. from 20 to 275 m deep, found on the bivalve mytilus adriaticus (fretter et al. 1986). subfossil finds. the skagen area, holocene, recorded from the atlantic and subatlantic. turbonilla crenata (brown 1827) distribution. from norway off the lofoten islands (thorson 1941) and south to the mediterranean. the species extends into the kattegat and øresund (fretter et al. 1986), but is absent from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from 15 to 350 m deep on fine sand. subfossil finds. the limfjord and north sea regions, holocene. recorded from the bælt sea and north sea during the eemian. fig. 35. odostomia umbilicaris (malm 1863). skagen 3, 82.34–82.50 m b.s., lab. no. 725,93. × 20. mguh 25345. geus bulletin no 3.pmd 28-06-2004, 08:4554 55 turbonilla delicata monterosato 1884 fig. 36 distribution. from the south-western part of the british isles to the mediterranean (fretter et al. 1986, p. 636). however, recorded from the northern kattegat and northern øresund although, rare (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. on soft bottoms, but at uncertain depths (fretter et al. 1986, p. 636). subfossil finds. the limfjord, north sea and skagen regions, holocene. from the skagen area recorded from the subatlantic. turbonilla lactea (linnaeus 1758) distribution. from northern norway and south to the mediterranean. present although uncommon in danish waters (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone to depths of about 80 m, occurring under stones in silty places in the tidal zone and on soft, muddy and sandy bottoms sublittorally (fretter et al. 1986, p. 634). subfossil finds. the limfjord, north sea and vendsyssel regions, holocene. recorded from the bælt sea, kattegat, and north sea regions during the eemian. subclass opisthobranchia order bullomorpha acteon tornatilis (linnaeus 1758) distribution. w and s iceland, norway off the lofoten islands and south to the mediterranean (lemche 1938). the species extends into the kattegat, including the øresund (petersen 1888, p. 78), but it is not recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone down to 250 m in sand, usually in sheltered areas (poppe & goto 1991, p. 192). subfossil finds. the limfjord, north sea and vendsyssel regions, holocene. recorded from the north sea and vendsyssel areas during the eemian. haminoea navicula (da costa 1778) distribution. from the british isles south to the mediterranean. occurrence. the lusitanian region. habitat. lives in the zostera beds in sheltered areas (poppe & goto 1991, p. 196). only subfossil finds. the bælt sea, kattegat, and north sea regions during the eemian. cylichna cylindracea (pennant 1777) distribution. w and s iceland, norway off the lofoten islands and south to the mediterranean. the species extends into the limfjord and kattegat, including øresund (petersen 1888, p. 78; jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from 40 to 200 m deep, in sand. according to lemche (1938) associated with the boreal amphiura filiformis community in deeper water south of iceland, corresponding to its occurrence in danish waters. subfossil finds. the limfjord and vendsyssel regions, holocene. fig. 36. turbonilla delicata (monterosato 1874). skagen 4, 20.0– 20.5 m b.s., lab. no. 345,93. × 20. mguh 25346. geus bulletin no 3.pmd 28-06-2004, 08:4555 56 cylichna alba (brown 1827) fig. 37 distribution.wand e greenland, around iceland, norway north of lofoten and south to the bay of biscay at greater depth (lemche 1938, p. 9). the species extends into the north sea and skagerrak (lemche 1928, p. 4). occurrence. the arctic, subarctic, boreal and lusitanian regions. habitat. according to lemche (1928), this species is also widely distributed in depth, being found in low water to depths down to 2700 m, on clay bottom (faeroes). subfossil finds. the limfjord, north sea and skagen areas, holocene. in the skagen well recorded from the subatlantic. also found in the vendsyssel area from the early/middle and late weichselian (the older and younger yoldia sea respectively). cylichna occulta (mighels 1841) distribution. e and w greenland, n and e iceland, norway north of lofoten (lemche 1938). occurrence. the arctic, subarctic and boreal (northernmost part) regions. will be referred as an arctic and subarctic species. habitat. from 10 to 388 m deep (iceland), especially found within the macoma calcarea community, and might also occur within the yoldia hyperborea community (lemche 1938). only subfossil finds. recorded from the kattegat region during the early/middle weichselian and the vendsyssel region during the early/middle and late weichselian (the older and younger yoldia sea respectively). scaphander lignarius (linnaeus 1758) distribution. from s and w iceland, the faeroes, norway off the lofoten islands and south to the mediterranean. occurrence. the boreal and lusitanian regions. habitat. from 60 to 700 m deep. the species is associated with the boreal spisula elliptica community on sandy plateaus south and west of iceland (lemche 1938, p. 7). subfossil finds. none. scaphander punctostriatus (mighels & adams 1841) distribution. e and w greenland, s and w iceland, the faeroes, norway north of lofoten and south to the mediterranean. occurrence. the arctic, subarctic, boreal and lusitanian regions. habitat. recorded from depths between 10 and 3000 m, this probably having some relation to its wide horizontal distribution. around the faeroes the species may be expected to be found on the great, sandy plateaus (lemche 1928, p. 4). subfossil finds. none. philine aperta (linnaeus 1767) distribution. from the faeroes, western norway off lofoten, and south to the mediterranean (lemche 1928). the species extends through kattegat and øresund, into the bælt sea as far as kieler bugt (petersen 1888, p. 83), and it is also recorded from the limfjord (petersen 1986a). fig. 37. cylichna alba (brown 1827). skagen 3, 44.88–45.00 m b.s., lab. no. 499,93. × 9.6. mguh 25347. geus bulletin no 3.pmd 28-06-2004, 08:4556 57 occurrence. the boreal and lusitanian regions. habitat. the species prefers shallow water but has been found at depths down to 100 m off the faeroes (lemche 1928) on sandy bottoms. in danish waters recorded from 10 to 30 m, also on sandy bottoms (petersen 1888). subfossil finds. the limfjord and vendsyssel regions, holocene. recorded from the bælt sea and the north sea during the eemian. philine angulata jeffreys 1867 this species is mentioned by jensen & knudsen (1995), but only on the basis of one broken shell found in the zoological museum in copenhagen; difficult to identify; occurrence uncertain. so in the light of “the difficult problem of the relation between punctata and angulata”, as treated by lemche (1948, p. 67), this find will be omitted, also that no subfossil species have been found. philine catena (montagu 1803) fig. 38 distribution. from norway off the lofoten islands and south to the mediterranean. occurrence. the boreal and lusitanian regions. habitat. from low water-mark to 76 m (forbes & hanley 1853) subfossil finds. the skagen region during the eemian. philine denticulata (adams 1800) distribution. from norway south to the mediterranean. occurrence. the boreal and lusitanian regions. subfossil finds. none. philine punctata (adams 1800) distribution. the faeroes, southern part of the west coast of norway, and south to the mediterranean. lemche (1928) also mentioned occurrences from greenland, which, however, was not repeated in later papers (lemche 1941a, b). jensen & knudsen (1995) report occurrences from the øresund, although rare, and petersen (1888) has a single find from the bælt sea and petersen (1986a) from the limfjord. occurrence. the boreal and lusitanian regions. habitat. at the faeroes the species is recorded from depths down to 240 m (lemche 1928). subfossil finds. the limfjord and vendsyssel regions, holocene. philine quadrata (wood 1839) distribution. w greenland, around iceland, norway north of lofoten, and south to the mediterranean. recorded from the øresund, but rare (jensen & knudsen 1995). occurrence. the subarctic, boreal and lusitanian regions. habitat. the vertical range of the species is about 35 m down to 2150 m (lemche 1938). subfossil finds. none. philine scabra (müller 1776) distribution. s and w iceland, the faeroes, norway north of lofoten, and south to the mediterranean. the species extends into the kattegat, including the øresund (petersen 1888, p. 84). occurrence. the boreal and lusitanian regions.fig. 38. philine cf. catena (montagu 1803). skagen 3, 180.77– 180.89 m b.s., lab. no. 798,93. × 40. mguh 25348. geus bulletin no 3.pmd 28-06-2004, 08:4557 58 habitat. off the coasts of w and s iceland from 20– 216 m on sandy bottom (lemche 1938, p. 11); however, in danish waters (kattegat) from 20–40 m on mixed bottom (petersen 1888, p. 84). subfossil finds. none. philinoglossa helgolandica hertling 1932 distribution. from the north sea – helgoland – and south to the mediterranean, the species might occur in danish waters (jensen & knudsen 1995, p. 29). occurrence. the boreal and lusitanian regions. habitat. probably in shell gravel (jensen & knudsen 1995). subfossil finds. none. order anaspidea diaphana minuta brown 1827 distribution. e and w greenland, around iceland, the faeroes, norway north of lofoten, and south to the mediterranean. the species extends into the kattegat and øresund (petersen 1888; jensen & knudsen 1995). occurrence. the arctic, subarctic, boreal and lusitanian regions. habitat. in general from the tidal zone to 770 m, but is said to prefer sandy clay at depths of 20–40 m (lemche 1928). this is very much the same as for the kattegat, where petersen (1888) says that the species prefers mixed bottom at a depth of about 19–38 m. subfossil finds. the limfjord and vendsyssel regions, holocene. retusa obtusa (montagu 1803) distribution. e and w greenland, around iceland, the faeroes, norway north of lofoten, and south to the british isles (shetland and scotland) (lemche 1928). the species is found in the limfjord and is common in the fjords and bays bordering the kattegat, including the øresund, and extends into the bælt sea and the western part of the baltic (petersen 1888, p. 81). occurrence. the arctic, subarctic and boreal regions. habitat. from the intertidal zone down to 300 m deep in mud or fine sand. in danish waters petersen (1888) points to the observed differences in depth, i.e. in the kattegat region around 20 m while 60 m deep in the baltic. the species is connected with the arctic macoma community (lemche 1941a, b). subfossil finds. the bælt sea, kattegat, limfjord, north sea and vendsyssel regions, holocene. also recorded from the vendsyssel area during the early/middle weichselian and the late weichselian (the older yoldia sea and the younger yoldia sea respectively). retusa truncatula (bruguière 1792) fig. 39 distribution. the faeroes, norway north of lofoten, and south to the mediterranean (lemche 1928, 1938). the species extends through the kattegat, including the øresund, and into the bælt sea, and it is also recorded from the limfjord (petersen 1888, p. 80). occurrence. the boreal and lusitanian regions. habitat. in general the species lives from the tidal zone down to 200 m (poppe & goto 1991). however, in danish waters, according to petersen (1888), it lives in shallow water down to only about 20 m deep on sandy bottoms with zostera. subfossil finds. the bælt sea, baltic, kattegat, limfjord, north sea, vendsyssel and skagen regions, holocene. fig. 39. retusa truncatula (bruguière 1792). skagen 3, 67.34–67.39 m b.s., core sample k-25. × 9.6. mguh 25349. geus bulletin no 3.pmd 28-06-2004, 08:4558 59 recorded from the skagen area from the subboreal and subatlantic. from the bælt sea and the north sea recorded during the eemian. retusa umbilicata (montagu 1803) fig. 40 distribution. from norway north of lofoten (lemche 1928) and south to the mediterranean (poppe & goto 1991). the species extends into the kattegat and øresund (petersen 1888; jensen & knudsen 1995), but is not recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from depths of about 20 to 30 m on mixed bottoms (petersen 1888). subfossil finds. the limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen well recorded from the subboreal and subatlantic. from the bælt sea,north sea, and vendsyssel regions found during the eemian. rhizorus acuminatus (bruguière 1792) distribution. from western and southern norway south to the mediterranean, in danish waters from the southern kattegat. occurrence. the boreal and lusitanian regions. habitat. from the tidal zone to 800 m deep. subfossil finds. none. akera bullata o.f. müller 1776 distribution. the faeroes, norway north of lofoten, and south to the mediterranean (lemche 1928). the species extends into the limfjord, the kattegat, including the øresund, and the bælt sea as far as the kieler bugt (petersen 1888). occurrence. the boreal and lusitanian regions. habitat. in sheltered bays down to 370 m deep (poppe & goto 1991). common in shallow water with zostera (petersen 1888). subfossil finds. the bælt sea, kattegat, limfjord and vendsyssel regions, holocene. in the bælt sea and kattegat regions recorded from the eemian. order thecosomata limacina retroversa (fleming 1823) fig. 41 distribution.eand w greenland, around iceland, norway north of lofoten, and south to ireland (lemche 1938). it is common in the north sea and skagerrak, penetrating into the kattegat, occasionally even into the bælt sea (kramp 1961). occurrence. the arctic, subarctic, boreal and lusitanian (northernmost) regions. habitat. pelagic. subfossil finds. recorded from the vendsyssel and skagen regions during the eemian and from the vendsyssel region during the late weichselian (the younger yoldia sea). fig. 40. retusa umbilicata (montagu 1803). skagen 3, 70.10–70.30 m b.s., lab. no. 721,95. × 20. mguh 25350. fig. 41. limacina retroversa (fleming 1823). skagen 3, 180.57– 180.70 m b.s., lab. no. 797,93. × 20. mguh 25351. geus bulletin no 3.pmd 28-06-2004, 08:4559 60 order gymnosomata clione limacina (phipps 1774) fig. 42 distribution. e and w greenland, around iceland, the faeroes, norway north of lofoten (lemche 1938), and south to the west coast of england. common in the northern part of the north sea and in the skagerrak, occasionally penetrating into the kattegat (kramp 1961). rare occurrences in the øresund (jensen & knudsen 1995). occurrence. the arctic, subarctic, boreal and lusitanian (northern part) regions. habitat. pelagic. subfossil finds. recorded (imprint) from the skagen well during the late weichselian (the younger yoldia sea). subclass pulmonata order basommatophora ovatella myosotis (draparnaud 1801) distribution. from scandinavia along the coast of western europe, south to the mediterranean (steenberg 1911, p. 204). in danish waters recorded from the bælt sea and the baltic (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. lives on sea wrack along the coasts (steenberg 1911). subfossil finds. none. lymnaea (radix) peregra (müller 1774) f. baltica linné distribution. lymnaea peregra is foundall overeurope: iceland, the faeroes, norway north of lofoten, and south to the mediterranean (mandahl-barth 1938). the species occurs in the baltic and southern øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. lymnaea peregra f. baltica linné and l. p. f. succinea nilsson are brackish water forms tolerating up to 8‰ salt (mandahl-barth 1949, p. 74). subfossil finds. the bælt sea and baltic regions, holocene. class scaphopoda order siphonodentalioida cadulus subfusiforme (m. sars 1865) fig. 43 distribution. western iceland, the faeroes, norway north of lofoten, and south to the mediterranean (knudsen 1949b). in danish waters taken in the skagerrak (muus 1959). occurrence. the boreal and lusitanian regions. habitat. from 80 to 1300 m deep on mud bottoms. in danish waters at a depth of 230 m on clay bottom (muus 1959). subfossil finds. recorded from the skagen well during the eemian. fig. 42. clione limacina (phipps 1774). skagen 3, 115.07–115.15 m b.s., lab. no. 749,93. × 20. mguh 25352. fig. 43. cadulus subfusiforme (m. sars 1865). skagen 3, 180.77– 180.89 m b.s., lab. no. 798,93. × 20. mguh 25353. geus bulletin no 3.pmd 28-06-2004, 08:4560 61 siphonodentalium lobatum (sowerby 1860) fig. 44 distribution. e and w greenland, n and e iceland (knudsen 1949b), the faeroes, norway north of lofoten, and south to portugal. the species might be found in norske rende. in norway often found in glacial deposits (muus 1959). occurrence. the main areas are the arctic, subarctic and boreal regions, but the species extends into the lusitanian region. habitat. from 36 to 3116 m on mud (knudsen 1949b). only subfossil finds. recorded from the skagen well in the late weichselian (the younger yoldia sea) and from the eemian. entalina tetragona (brocchi 1814) fig. 45 distribution. from northern norway south to the bay of biscay and the mediterranean. from danish waters in the north sea and skagerrak (muus 1959). occurrence. the boreal and lusitanian regions. habitat. it is a deep-water species preferring mud bottoms (poppe & goto 1993). in the danish waters known from 100 to 480 m on mixed bottom (muus 1959) connected with the amphilipis norvegica/pecten vitreus community at depths of 250–700 m. subfossil finds. recorded from the skagen well during the eemian. order dentalioida antalis agile g.o. sars 1878 distribution. s and w iceland, norway off the lofoten islands, and south to the mediterranean. in the danish waters the species extends from the north sea into the skagerrak and kattegat. occurrence. the boreal and lusitanian regions. habitat. from 55 m down to 1250 m. in scandinavian waters rarely at depths less than 70 m (muus 1959). subfossil finds. none. antalis entalis (linnaeus 1758) distribution. w greenland, iceland, the faeroes, norway north of lofoten, and south to the mediterranean (knudsen 1949b). common in the danish waters, extending into the kattegat (muus 1959) although rare in the øresund (jensen & knudsen 1995). occurrence. the subarctic, boreal and lusitanian regions. habitat. from 20 to 400 m in the kattegat region, and in the north sea between 30 and 200 m (muus 1959). however, the vertical range in general goes from 6– 3200 m (knudsen 1949b). subfossil finds. the skagen well during the eemian. fig. 44. siphonodentalium lobatum (sowerby 1860). skagen 3, 114.64–? m b.s., core sample k-57. × 9.6. mguh 25354. fig. 45. entalina tetragona (brocchi 1814). skagen 3, 183.77– 184.00 m b.s., lab. no. 799,93. × 9.6. mguh 25355. geus bulletin no 3.pmd 28-06-2004, 08:4561 62 dentalium vulgare da costa 1778 fig. 46 distribution. from the british isles and south to the mediterranean. might be found in the southern north sea. occurrence. mainly the lusitanian region. habitat. sublittorally from 1 m and down to 50 m on mud and sandy bottoms. only subfossil finds. recorded from the skagen well during the eemian. class bivalvia subclass palaeotaxodonta order nuculoida nucula nitidosa winckworth 1930 fig. 47 distribution. from norway off the lofoten islands and south to the mediterranean (madsen 1949). the species extends from the more shallow part of the north sea and skagerrak into the kattegat, including the øresund. very common in the western part of the limfjord (jensen & spärck 1934, p. 23). occurrence. the boreal and lusitanian regions. habitat. in general from 7 to 250 m deep on fine sand or sand/mud bottoms (poppe & goto 1993). however, according to jensen & spärck (1934) the species is often found at depths of 6–10 m in the limfjord and 10–30 m in the kattegat region. subfossil finds. the limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen well recorded from the subatlantic. during the eemian found in the bælt sea, baltic and north sea regions. nucula nucleus (linnaeus 1767) fig. 48 distribution. from norway off the lofoten islands, the faeroes, and south to the mediterranean (madsen 1949). posselt & jensen (1898) have but few records from west greenland. the species extends from the north sea into the kattegat, including the øresund (jensen & spärck 1934), and is recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. lives from the tidal zone down to 150 m on gravel and mud bottoms. in danish waters common between 20 and 100 m. subfossil finds. the limfjord, north sea, vendsyssel fig. 46. dentalium vulgare da costa 1778. skagen 3, 181.17– 181.21 m b.s., core sample k-99. × 9.6. mguh 25356. fig. 47. nucula nitidosa winckworth 1930. skagen 4, 21.0–21.5 m b.s., lab. no. 346,93. × 9.6. right valve. mguh 25357. fig. 48. nucula nucleus (linnaeus 1767). skagen 4, 9.0–9.5 m b.s., lab. no. 334,93. × 4.8. right valve. mguh 25358. geus bulletin no 3.pmd 28-06-2004, 08:4562 63 and skagen regions, holocene. recorded in the skagen well from the subatlantic. in the eemian recorded from the baltic and vendsyssel regions. nucula sulcata (bronn 1831) distribution. norway off the lofoten islands and south to the mediterranean. in danish waters common in the deeper parts of the kattegat and øresund (jensen & knudsen 1995), and it is also recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from 10 to 400 m deep, on mud or clay bottoms, and down to 2250 m (poppe & goto 1993). in kattegat from 50 to 100 m deep on silty bottom (jensen & spärck 1934). subfossil finds. the north sea, holocene. from the bælt sea, north sea and vendsyssel regions recorded from the eemian. nuculoma hanleyi winckworth 1931 distribution. from the british isles south to spain (poppe & goto 1993). recorded from the northern kattegat (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from 30 to 90 m deep on mud and gravel bottoms. subfossil finds. none. nuculoma tenuis (montagu 1808) distribution. w and e greenland, spitsbergen, around iceland, the faeroes, norway north of lofoten, and south to the mediterranean (madsen 1949). the species extends from the deeper part of the skagerrak into the kattegat, øresund, and bælt sea (jensen & spärck 1934). occurrence. the arctic, subarctic, boreal and lusitanian regions. habitat. from off-shore down to 300 m on muddy bottoms (poppe & goto 1993). the species is more littoral in the northern latitudes than in the south (jensen & spärck 1934). subfossil finds. from the limfjord and north sea regions, holocene. recorded from the vendsyssel area during the eemian, and from the kattegat and vendsyssel area during the early/middle weichselian (the older yoldia sea), and furthermore from the vendsyssel region during the late weichselian (the younger yoldia sea). nuculana minuta (müller 1776) fig. 49 distribution. se and w greenland, around iceland, the faeroes, norway north of lofoten, and south to the british isles. according to ockelmann (1958), lacking in the most high-arctic seas. in south-western europe only at depths greater than 400 m (madsen 1949). the species extends from the north sea and skagerrak into the kattegat and øresund (jensen & spärck 1934). occurrence. the subarctic and boreal regions. habitat. from 10 to 190 m. however, recorded from 2000 m on mud, sand and gravel bottoms (poppe & goto 1993). in danish waters common on silty bottom (jensen & spärck 1934). subfossil finds. the skagen region, holocene. recorded from the skagen well during the subboreal and subatlantic. from the vendsyssel region found during the eemian and in the skagen area during the late weichselian (the younger yoldia sea). fig. 49. nuculana minuta (müller 1776). geus collection. læsø, denmark. × 4.8. right valve. mguh 25359. geus bulletin no 3.pmd 28-06-2004, 08:4563 64 nuculana pernula (müller 1776) fig. 50 distribution. e and w greenland, around iceland, the faeroes,norwaynorthof lofoten, and south to the bay of biscay. however, in sw europe only at depths greater than 400 m (madsen 1949). in danish waters the species extends from the north sea, skagerrak, and into the kattegat, øresund, and bælt sea region north of femern (jensen & spärck 1934). occurrence. the arctic, subarctic and boreal regions. habitat. lives off-shore between 80 and 900 m deep, especially in mud bottoms (poppe & goto 1993). however, in danish waters the species is found from a depth of 20 m (kattegat) to 200 m (skagerrak) according to jensen & spärck (1934), and in the arctic (east greenland) from 3–9 m (ockelmann 1958), being mostly littoral in the arctic. subfossil finds. recorded from the vendsyssel and skagen regions during the eemian. from the kattegat, vendsyssel and skagen regions during the early and middle weichselian (the older yoldia sea). from the vendsyssel and skagen regions during the late weichselian (the younger yoldia sea). yoldia hyperborea lovén 1859 fig. 51 distribution. e and w greenland, spitsbergen, around iceland and norway north of lofoten (madsen 1949). occurrence. the arctic, subarctic and boreal (northern part – high-boreal) regions. habitat. from about 5 to 675 m on clay or mud, in few cases sand (madsen 1949). only subfossil finds. recorded from the vendsyssel and skagen areas during the early and middle weichselian (the older yoldia sea) and from the skagen well also during the late weichselian (the younger yoldia sea). portlandia arctica (gray 1824) fig. 52 distribution. according to ockelmann (1958, p. 26): “widely distributed in high-arctic seas”. occurrence. north greenland, east greenland, spitsbergen, the barents sea, novaya zemlya, the kara sea, the siberian ice sea, viz: arctic and subarctic regions. habitat. from 2 m to 340 m deep, however, the species is most common at depths between 10 and 50 m on a muddy or clayey bottom (ockelmann 1958, p. 25). fig. 50. nuculana pernula (müller 1776). geus collection. læsø, denmark. × 4.8. right valve. mguh 25360. fig. 51. yoldia hyperborea lovén 1859. geus collection. nordre strømfjord, greenland. × 4.8. right valve. mguh 25361. fig. 52. portlandia arctica (gray 1824). geus collection. disko, greenland, 46 m a.s.l. × 4.8. specimen seen from the right. mguh 25362. geus bulletin no 3.pmd 28-06-2004, 08:4564 65 only subfossil finds. recorded from the kattegat, vendsyssel, and skagen areas during the early and middle weichselian (the older yoldia sea) and from the same areas during the late weichselian (the younger yoldia sea). yoldiella lucida (lovén 1846) distribution. w greenland, around iceland, norway north of lofoten (madsen 1949), and south over the british isles to the mediterranean. in the danish waters found in the deeper part of the skagerrak (jensen & spärck 1934). occurrence. the subarctic, boreal and lusitanian regions. habitat. from about 20 m to about 1400 m (iceland), in the skagerrak from about 200 m to more than 600 m on clayey bottoms. subfossil finds. from the vendsyssel area during the eemian. yoldiella lenticula (möller 1842) fig. 53 distribution. e and w greenland, spitsbergen, siberian sea, norway north of the lofoten islands, south to the british isles. further south only at depths greater than 400 m (madsen 1949). occurrence. the arctic, subarctic and boreal regions. habitat. most common at depths between 20 and 200 m (east greenland) on rather pure clay or mud (ockelmann 1958). only subfossil finds. recorded from the vendsyssel area during the eemian, early/middle weichselian (the older yoldia sea) and late weichselian (the younger yoldia sea). found in the skagen area from the late weichselian (the younger yoldia sea). yoldiella frigida (torell 1859) fig. 54 distribution. e and w greenland, spitsbergen, around iceland, norway from north of lofoten, and south to the mediterranean (madsen 1949). in the danish waters found in the deeper parts of the skagerrak (jensen & spärck 1934). however, this is questioned by ockelmann (1958), who says that the main distribution is high-arctic and therefore the recent finds in danish waters should be referable to yoldiella nana. however, according to the old information, the species must have a wide distribution. occurrence. the arctic and subarctic regions, boreal and lusitanian. habitat. most common at depths between 30 and 150 m on bottoms consisting of clay, mud, and clay mixed with sand and gravel (ockelmann 1958). subfossil finds. following the distribution of portlandia frigida sensu jensen & spärck (1934), the species has been recorded from the vendsyssel and skagen areas during the eemian and from the early and middle weichselian (the older yoldia sea) and the late weichselian (the younger yoldia sea). fig. 53. yoldiella lenticula (möller 1842). skagen 3, 121.39– 121.50 m b.s., lab. no. 759,93. × 20. specimen seen from the left. mguh 25363. fig. 54. yoldiella frigida (torell 1859). skagen 3, 127.39–127.50 m b.s., lab. no. 763,93. × 20. specimen seen from the right. mguh 25364. geus bulletin no 3.pmd 28-06-2004, 08:4565 66 yoldiella philippiana (nyst 1845) distribution. from norway off the lofoten islands and south to the mediterranean. occurrence. the boreal and lusitanian regions. habitat. lives to a depth of about 135 m (poppe & goto 1993); however, found at depths of 2500 m (jensen & spärck 1934). only subfossil finds. recorded from the vendsyssel area during the eemian. yoldiella nana (m. sars 1846) distribution. considering the discussion by ockelmann (1958) on distinguishing between portlandia frigida and yoldiella nana, where the latter “at least in part” should be referable to p. fraterna, it is not possible to give any information on the distribution of yoldiella nana mentioned by jensen & knudsen (1995) as being part of the recent danish fauna. subfossil finds. none. malletia obtusa (g.o. sars 1872) distribution. from norway off the lofoten islands and south to sw europe. the occurrences around the british isles and in the mediterranean are at depths greater than 400 m (madsen 1949). occurrence. the boreal and lusitanian regions. habitat. this species has a wide range of depths according to madsen (op. cit.) from 20–3200 m deep. in danish waters only from the deeper part (> 300 m), and in the skagerrak mostly at a depth of 600 m (jensen & spärck 1934). subfossil finds. none. subclass pteriomorphia order arcoida acar nodulosa (müller 1766) distribution. s and w iceland, the faeroes, norway north of lofoten, and south to the mediterranean. the species has been taken on the dogger bank, but not in the inner danish waters. occurrence. the boreal and lusitanian regions. habitat. from the tidal zone down to 1000 m fixed with its byssus to hard substrates (poppe & goto 1993). subfossil finds. none. bathyarca glacialis (gray 1824) fig. 55 distribution. w and e greenland, spitsbergen, and around iceland. the occurrences from sw europe are of dead shells and from deep water (ockelmann 1958). occurrence. mainly the arctic and subarctic regions; however, the occurrences from southern iceland imply extension into the boreal region as well. habitat. from 6–10 m down to 425 m (east greenland), but most abundant below 40 m on clay bottoms with stones and gravel, where the astarte crenata community occurs. only subfossil finds. recorded from the vendsyssel (not in the skærumhede sequence) and skagen regions during the early/middleweichselian and the late weichselian respectively. bathyarca pectunculoides (scacchi 1834) distribution.wand e greenland, around iceland, norway north of lofoten, and south to the mediterranean. in the danish waters the species is rather common in the deeper part of the skagerrak between 300–700 m (jensen & spärck 1934). fig. 55. bathyarca glacialis (gray 1824). skagen 3, 114.64–? m b.s., core sample-57. × 9.6. fragments of left valve. mguh 25365. geus bulletin no 3.pmd 28-06-2004, 08:4566 67 occurrence. the arctic, subarctic, boreal and lusitanian regions. habitat. from 50 m (the shetland isles) to more than 2000 m (the mediterranean) (madsen 1949). subfossil finds. none. order mytiloida mytilaster lineatus (gmelin 1791) distribution. a mediterranean species according to jensen & spärck (1934) but also found in the neighbouring atlantic (poppe & goto 1993). occurrence. the lusitanian region. habitat. intertidal, attached to rocks. only subfossil finds. the eemian in the kattegat and north sea regions. mytilaster solidus form minimus (poli 1795) distribution. from bretagne and south into the mediterranean. occurrence. the lusitanian region. habitat. intertidal, attached to rocks or algae. subfossil finds. recorded from the bælt sea during the eemian. mytilus edulis linnaeus 1758 fig. 56 distribution. w and e greenland, but only along the south-eastern coast (ockelmann 1958), around iceland, the faeroes, norway north of lofoten, and south to the mediterranean. known from all parts of the danish waters, including the baltic. occurrence. the subarctic, boreal and lusitanian regions. habitat. intertidal to 40 m deep, but in danish waters common as an epifaunal element down to a depth of 10 m; however, in the baltic as deep as 40 m (jensen & spärck 1934). subfossil finds. the bælt sea, baltic, kattegat, limfjord, north sea, vendsyssel and skagen regions, holocene. recorded in the subatlantic from the skagen well. in the eemian records from the bælt sea, baltic, kattegat, north sea and vendsyssel regions. in the late weichselian also recorded from the vendsyssel region (the younger yoldia sea). modiolula phaseolina (philippi 1844) distribution. from sw and nw iceland, the faeroes, norway north of lofoten, and south to the british isles (petersen 1968) and the mediterranean. the species extends from the north sea and skagerrak into the kattegat and øresund (jensen & knudsen 1995), but is not recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone and down to 160 m, attached by its byssus to rocks or on the base of the larger seaweeds (poppe & goto 1993), but also recorded from depths of 1000 m. subfossil finds. the bælt sea, limfjord and vendsyssel regions, holocene. recorded from the bælt sea and the north sea regions during the eemian. modiolus adriaticus (lamarck 1819) distribution. recorded from the southern part of the british isles and denmark (petersen 1968) south to the mediterranean. in danish waters taken in the kattegat, including the øresund (jensen & spärck 1934), but fig. 56. mytilus edulis linnaeus 1758. skagen 4, 18.7–18.8 m b.s., lab. no. 311,95. × 4.8. fragment of left valve. mguh 25366. geus bulletin no 3.pmd 28-06-2004, 08:4567 68 not observed in the limfjord (petersen 1986a). it might have been passed over in many places, as mentioned by petersen (1888, p. 127). occurrence. the boreal and lusitanian regions. habitat. sublittoral between 14 and 75 m on mud bottoms. subfossil finds. the limfjord and vendsyssel region, holocene. modiolus modiolus (linnaeus 1758) distribution. around iceland, the faeroes, norway north of lofoten, and south to the british isles (petersen 1968) and the bay of biscay (poppe & goto 1993). in danish waters, including the limfjord, the species extends into the bælt sea (jensen & spärck 1934) and the øresund (jensen & knudsen 1995). occurrence. the subarctic, boreal and lusitanian regions. habitat. from the extreme low tide down to 150 m attached with its byssus to rocks or gravel (poppe & goto 1993). in danish waters the species replaces mytilus edulis as the dominating epifaunal element in deeper water (jensen & spärck 1934). subfossil finds. the bælt sea, limfjord and vendsyssel regions, holocene. recorded from the baltic and north sea during the eemian. musculus discors (linnaeus 1767) fig. 57 distribution. e and w greenland, around iceland and spitsbergen (madsen 1949), norway from north of lofoten (petersen 1968) and south to the mediterranean. in danish waters the species extends into the bælt sea and øresund (jensen & spärck 1934), but it has not been observed in the limfjord (petersen 1986a). occurrence. the arctic, subarctic, boreal, and lusitanian regions. habitat. from the intertidal zone on algae (poppe & goto 1993) and rarely on water deeper than about 200 m (jensen & spärck 1934). subfossil finds. the bælt sea, kattegat, limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen well recorded from the subatlantic. musculus laevigatus (gray 1824) distribution. e and w greenland, around iceland and norway north of lofoten. occurrence. the arctic, subarctic and boreal regions. habitat. from the infralittoral zone down to 83 m (poppe & goto 1993). only subfossil finds. recorded from the early/middle weichselian (the older yoldia sea, but not in the skærumhede sequence) and the late weichselian (the younger yoldia sea) in the vendsyssel region. musculus niger (gray 1824) distribution. e and w greenland, around iceland, the faeroes, norway north of lofoten, and south to the north sea and the irish sea (madsen 1949; petersen 1968). in danish waters the species extends from the north sea into the kattegat, øresund and bælt sea as far as warnemünde (jensen & spärck 1934; jensen & knudsen 1995). occurrence. the arctic, subarctic and boreal regions. habitat. most often found in danish waters at water depths of more than 25 m (jensen & spärck 1934). poppe & goto (1993) indicate from 7 m deep down to about 135 m. however, off the east greenland coast the species is rarely met with at depths exceeding 40 m (ockelmann 1958). fig. 57. musculus discors (linnaeus 1767). skagen 4, 28.0–28.5 m b.s., lab. no. 353,93. × 20. left valve. mguh 25367. geus bulletin no 3.pmd 28-06-2004, 08:4568 69 subfossil finds. recorded from the vendsyssel region during the eemian, the early and middle weichselian (the older yoldia sea) and the late weichselian (the younger yoldia sea). modiolaria tumida (hanley 1843) distribution. the british isles and the shetlands, and south to the mediterranean (madsen 1949). in danish waters the species is very common in the limfjord, but also in the other fjords, and it extends into the bælt sea (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. it is a common shallow-water species in danish waters (jensen & spärck 1934), but goes down to 60 m (poppe & goto 1993) associated with tunicates and echinoderms. subfossil finds. the limfjord and vendsyssel regions, holocene. recorded from the north sea region during the eemian. crenella decussata (montagu 1803) distribution.eand w greenland, around iceland, norwaynorthof lofoten and south to the british isles (madsen 1949; petersen 1968). in danish waters the species is found in the kattegat, including the øresund. occurrence. the arctic, subarctic and boreal regions. habitat. from 4 to 200 m deep on all kinds of bottoms (poppe & goto 1993). according to jensen & spärck (1934), most common in danish waters between 15 and 30 m. subfossil finds. recorded in the kattegat during the holocene and from the early and middle weichselian (the older yoldia sea) in the vendsyssel region. adipicola simpsoni (marshall 1900) distribution. from southern iceland and south to portugal and the mediterranean. might be found in danish waters (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. as mentioned by poppe & goto (1993, p. 48): “the species has been collected repeatedly on the skulls of whales, where it lies, attached in the sutures, by its byssus”. subfossil finds. none. order pteroida chlamys islandica (o.f. müller 1776) distribution.seandwgreenland, around iceland, spitsbergen, norway north of lofoten, and south to the shetlands and the orkney islands (rare) (petersen 1968). occurrence. the subarctic and boreal regions. habitat. around iceland one of the most common bivalves present from nearly all localities along the nw, n and e coast, both in the fjords and on the outer part of the shelf, at depth from a few metres to 300 m (madsen 1949). common in danish waters from 10 to 100 m (jensen & spärck 1934) on rocks and gravel bottoms. lives attached to hard substrates with its byssus (poppe & goto 1993). only subfossil finds. recorded from the vendsyssel region during the early and middle weichselian (the olderyoldia sea) and the late weichselian (the younger yoldia sea). aequipecten opercularis (linnaeus 1758) distribution. norway from north of lofoten, and south to the mediterranean (petersen 1968). in danish waters the species extends into the kattegat and øresund, but is not recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone down to 400 m on all types of bottoms except rocky ones (poppe & goto 1993). subfossil finds. the limfjord and vendsyssel regions, holocene. recorded from the north sea region during the eemian. geus bulletin no 3.pmd 28-06-2004, 08:4569 70 chlamys varia (linnaeus 1758) fig. 58 distribution. norway off the lofoten islands, and south to the mediterranean (petersen 1968). in danish waters the species occurs in the limfjord (jensen & spärck 1934) and has been recorded juvenile from the øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. intertidal to 83 m, attached by its byssus (poppe & goto 1993). subfossil finds. the limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen well recorded from the subatlantic. recorded from the north sea during the eemian. delectopecten vitreus (gmelin 1791) fig. 59 distribution. w greenland, w iceland, spitsbergen, norway north of lofoten, and south to the british isles (madsen 1949), and according to poppe & goto (1993) also into the mediterranean. in danish waters from the skagerrak (jensen & spärck 1934). occurrence. the subarctic, boreal and lusitanian regions. habitat. between 30 and 600 m, fixed by its byssus to hard substrates (poppe & goto 1993). the species lives in the deeper part of the skagerrak, from 400 to 600 m, according to jensen & spärck (1934). subfossil finds. the limfjord, holocene. recorded from the eemian in the skagen well. palliolum greenlandicum (sowerby 1842) fig. 60 distribution. e and w greenland, n and e iceland, spitsbergen, and norway north of lofoten (madsen 1949). however, at depths greater than 400 m the species has been found off the faeroes and the british isles. occurrence. the arctic, subarctic and boreal regions. habitat. in the arctic seas living in shallow water from 5 m, but most common between 20 and 70 m on clay bottoms containing stones or shells (ockelmann 1958). only subfossil finds. recorded from the vendsyssel region and the skagen well during the early/middle weichselian (the older yoldia sea). fig. 58. chlamys varia (linnaeus 1758). skagen 3, 32.85–32.90 m b.s., core sample-2. × 20. right valve. mguh 25368. fig. 59. delectopecten cf. vitreus (gmelin 1791). skagen 3, 183.17–183.40 m b.s., lab. no. 9e+05. × 9.6. mguh 25369. fig. 60. palliolum greenlandicum (sowerby 1842). geus collection. east greenland. × 4.8. right valve. mguh 25370. geus bulletin no 3.pmd 28-06-2004, 08:4570 71 palliolum striatum (müller 1776) distribution. from s and w iceland, the faeroes, norway north of lofoten, and south to the mediterranean. the species extends from the north sea and skagerrak into the kattegat, including øresund. occurrence. the boreal and lusitanian regions. habitat. from shallow water around 5 m to more than 800 m deep on all types of bottom (poppe & goto 1993). the vertical range off iceland is indicated to lie between 100 and 260 m (madsen 1949). subfossil finds. the limfjord region, holocene. palliolum tigerinum (müller 1776) distribution. from nw, w and s iceland, the faeroes, norway north of lofoten, and south to morocco. in danish waters from the north sea and skagerrak into the kattegat, including the øresund. occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone down to 400 m, but deeper in the southern part of its range on sandy bottoms (poppe & goto 1993). subfossil finds. the limfjord region, holocene. pecten maximus (linnaeus 1758) distribution. norway off the lofoten islands and south to spain. in danish waters rarely found living in the kattegat and only shells have been recovered from the øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone down to 250 m on sand and gravel bottoms (poppe & goto 1993). subfossil finds. recorded from the vendsyssel region, holocene. pseudamussium septemradiatum (müller 1776) distribution. s iceland, norway from north of the lofoten islands, and south to the mediterranean. in danish waters the species extends from the skagerrak and becomes common in the southern kattegat with finds also in the øresund (jensen & spärck 1934; jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. in general living between 60 and 600 m deep on muddy bottoms (poppe & goto 1993). in danish waters often found between 30 and 60 m. subfossil finds. recorded from the vendsyssel region during the eemian. similipecten similis (laskey 1811) distribution. s and w iceland, the faeroes, norway north of lofoten, and south to the mediterranean (petersen 1968). the species extends into the kattegat from skagerrak. occurrence. the boreal and lusitanian regions. habitat. sublittorally between 4 and 250 m deep on sand and fine gravel bottoms (poppe & goto 1993). in danish waters from 30 to 80 m deep (jensen & spärck 1934). however, the vertical range around iceland is 200–320 m (madsen 1949). subfossil finds. recorded from the vendsyssel area during the eemian. pododesmus patelliformis (linnaeus 1761) distribution. s and w iceland, the faeroes, norway off the lofoten islands, and south to the mediterranean (madsen 1949). the species extends from the north sea into the kattegat, including the øresund, and occurs also in the limfjord (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. intertidal to 50 m deep on gravel or rock bottoms, often attached to shells (poppe & goto 1993). subfossil finds. the limfjord and vendsyssel regions, holocene. pododesmus squama (gmelin 1791) distribution. around the british isles and from danish waters the occurrences in the kattegat and øresund geus bulletin no 3.pmd 28-06-2004, 08:4571 72 are questioned, although larvae occur (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone to 75 m deep on all types of bottoms attached to hard substrates (poppe & goto 1993). subfossil finds. none. anomia ephippium linnaeus 1758 distribution. from the british isles, including the orkney islands, and south to the mediterranean. occurrence. mainly the lusitanian region. habitat. from the intertidal zone down to 150 m on all kinds of hard substrates. only subfossil finds. the limfjord and the vendsyssel region, holocene. heteranomia squamula (linnaeus 1758) fig. 61 distribution. around iceland, the faeroes, norway north of lofoten, and south to the bay of biscay. the species extends into the kattegat, øresund, and occurs in the limfjord (jensen & spärck 1934). occurrence. the subarctic, boreal and lusitanian regions. habitat. from 5 to 110 m deep fixed on hard substrates, but also on algae and crustaceans (poppe & goto 1993). subfossil finds. the kattegat, limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen well recorded from the subatlantic. during the eemian recorded from the bælt sea and the north sea regions. crassostrea gigas (gmelin 1791) this oyster species from the portuguese–spanish region (poppe & goto 1993) has been introduced in 1972 as spat for commercial production (jensen & knudsen 1995) and is not considered here, although mentioned as now part of the danish molluscan fauna. no subfossil records either. ostrea edulis linnaeus 1758 fig. 62a, b distribution. from the southern part of the west coast of norway south to the mediterranean. in danish waters only common in the western part of the limfjord, although stray specimens are found in the northern north sea, skagerrak and northern kattegat (jensen & spärck 1934; jensen & knudsen 1995). fig. 61. heteranomia squamula (linnaeus 1758). skagen 4, 21.0– 21.5 m b.s., lab. no. 346,93. × 20. left valve. mguh 25371. fig. 62. a, b: ostrea edulis linnaeus 1758. skagen 4, 8.0–8.5 m b.s., lab. no. 333,93. × 9.6. left valve of juvenile specimen (exterior and interior, respectively). mguh 25372. geus bulletin no 3.pmd 28-06-2004, 08:4572 73 occurrence. the boreal and lusitanian regions. habitat. from the tidal zone down to 90 m on all types of bottoms. in danish waters the species can be found at depths of 3 to 7 m in the limfjord, but also deeper elsewhere (jensen & spärck 1934). subfossil finds. the bælt sea, kattegat, limfjord, north sea, vendsyssel and skagen regions, holocene. recorded from the subatlantic in the skagen well. finds from the bælt sea, kattegat, and north sea regions during the eemian. limaria hians (gmelin 1791) distribution. norway north of lofoten, the orkney islands and south to the mediterranean (petersen 1968). a few records from the northern and central parts of the kattegat (jensen & knudsen 1995). already jensen & spärck (1934) mentioned that the species then known from the deeper parts of the north sea might occur in danish waters. occurrence. the boreal and lusitanian regions. habitat. lives from the low tide mark zone down to 100 m on coarse sand and gravel bottoms (poppe & goto 1993). subfossil finds. none. limaria loscombi (sowerby 1832) distribution. norway off the lofoten islands, the faeroes and south to the mediterranean (petersen 1968). poppe & goto (1993) mentioned occurrences of l. hians and l. loscombi off iceland which, however, cannot be found in the other literature. in danish waters the species has been found in the kattegat, including the øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. lives from 35 to 100 m deep on fine sand and sand–mud bottoms (poppe & goto 1993). in danish waters it is characteristic in the south-eastern part of kattegat together with pseudamussium septemradiatum (jensen & spärck 1934). subfossil finds. none. limatula subauriculata (montagu 1808) distribution. se and w greenland, n and e iceland, norway north of lofoten, and south to the mediterranean (petersen 1968). in the danish waters only shells have been recorded from the northern kattegat (jensen & knudsen 1995). occurrence. the subarctic, boreal and lusitanian regions. habitat. especially living on the continental shelves. however, records ranges from 4 to 2000 m (poppe & goto 1993). subfossil finds. none. subclass heterodonta order veneroida chama gryphoides linnaeus 1767 this lusitanian species (up to the coasts of portugal) has only one record from danish waters (jensen & knudsen 1995), and this is considered to have been dropped by a ship. therefore it will not be discussed. no subfossil records. lucinella divaricata (linnaeus 1758) distribution. from the english channel and southern part of the north sea south to the mediterranean. occurrence. mainly the lusitanian region. habitat. from the tidal zone down to a depth of 60 m in fine sand and/or mud (poppe & goto 1993). only subfossil finds. the eemian in the bælt sea, kattegat, and north sea regions. lucinoma borealis (linnaeus 1758) distribution. the faeroe islands, norway from off lofoten, and south to the mediterranean (petersen 1968). in danish waters the species occurs in the kattegat, including the øresund, but is not recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. geus bulletin no 3.pmd 28-06-2004, 08:4573 74 habitat. from the intertidal zone down to 500 m deep on gravel bottoms and in pure sand and/or mud (poppe & goto 1993). in danish waters between 20 and 50 m (jensen & spärck 1934). subfossil finds. the limfjord and vendsyssel regions, holocene. myrtea spinifera (montagu 1803) distribution. from norway south to morocco. in danish waters recorded from the northern kattegat (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from 7 to 250 m deep on sand, mud and gravel bottoms (poppe & goto 1993). subfossil finds. none. axinopsida orbiculata (g.o. sars 1878) distribution. e and w greenland, around iceland, the faeroes, and norway north and just south of the lofoten islands (petersen 1968). the species occurs off the north western part of scotland. occurrence. the arctic, subarctic and boreal regions. habitat. from 2 to 50 m deep on sand, clay and mud around iceland (madsen 1949). north of the hebrides occurring at depths down to 900 m (jensen & spärck 1934). only subfossil finds. the vendsyssel region from the early and middle weichselian (the older yoldia sea) and the late weichselian (the younger yoldia sea). thyasira croulinensis (jeffreys 1847) distribution. w greenland, around iceland, the faeroes, norway north of lofoten (petersen 1968), and south to the mediterranean (poppe & goto 1993). in danish waters taken in the north sea and skagerrak. occurrence. the subarctic, boreal and lusitanian regions. habitat. about 40 to 2500 m off the faeroes on gravel and clay (petersen 1968). in danish waters the species is found in the deeper water (jensen & spärck 1934). subfossil finds. none. thyasira equalis (verrill & bush 1898) distribution. it is questioned by nordsieck (1969, p. 79) if t. equalis should be thyasira flexuosa var. rotunda. there are no subfossil finds under the name of t. equalis, so this species will not be considered any further. however, as discussed by ockelmann (1958, p. 100) a species, t. equalis, does occur in the arctic, while t. flexuosa has a boreo-lusitanian main distribution. thyasira flexuosa (montagu 1803) fig. 63 distribution. e and w greenland, spitsbergen, around iceland, the faeroes, norway north of lofoten, and south to the mediterranean. in the danish waters it is very common and extends into the øresund (jensen & knudsen 1995), but is not recorded from the limfjord (petersen 1986a). occurrence. the arctic, subarctic, boreal and lusitanian regions (see comments under t. equalis). habitat. from 10 to 2000 m deep on sand and mud bottoms (poppe & goto 1993). in danish waters from 20 m to around 100 m deep on clay bottoms (jensen & spärck 1934). from the north sea recorded at 30 to 200 m depths on mixed bottom in the trenches around the dogger bank (petersen 1977). subfossil finds. the limfjord, north sea, vendsyssel fig. 63. thyasira flexuosa (montagu 1803). skagen 3, 35.90– 36.00 m b.s., lab. no. 496,93. × 9.6. left valve. mguh 25373. geus bulletin no 3.pmd 28-06-2004, 08:4574 75 and skagen regions, holocene. from the skagen well recorded from the subatlantic. from the north sea during the eemian and in the vendsyssel region recorded from the late weichselian (the younger yoldia sea). thyasira sarsi (philippi 1845) distribution. from novaja semlja along the coast of norway south to the skagerrak region. the species extends into the kattegat, including the øresund (jensen & knudsen 1995). occurrence. the arctic, subarctic and boreal regions. habitat. from 100 m to deep water. subfossil finds. none. leptaxinus ferruginosus (forbes 1844) distribution. w greenland, w iceland, spitsbergen, norway off the lofoten islands, and southwards to madeira (madsen 1949). jensen & spärck (1934) mentioned the species from the deeper part of the skagerrak,but jensen & knudsen (1995) have no further record of this species as belonging to the recent danish fauna. occurrence. the arctic, subarctic, boreal and lusitanian regions. habitat. about 20 m to more than 3000 m. the vertical range off iceland is 320–560 m. so here is another example of tropical submerge (see order mesogastropoda natica affinis). subfossil finds. recorded from the vendsyssel region during the eemian. mysella bidentata (montagu 1803) fig. 64 distribution. around iceland, the faeroes, norway from north of lofoten, and south to west africa (petersen 1968). the species has a common distribution in danish waters, including the limfjord, south to the bælt sea by warnemünde (arntz et al. 1976), and the øresund (jensen & spärck 1934). also found in the north sea (petersen 1977). occurrence. the subarctic, boreal and lusitanian regions. habitat. intertidal zone down to 2500 m, often in commensal association with other animals. in the north sea it is recorded from 20 to 90 m as a commensal on for example, acrocnida brachiata (petersen 1977). subfossil finds. the bælt sea, kattegat, limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen well, records from the preboreal–boreal, subboreal and subatlantic. from the bælt sea, baltic, north sea and vendsyssel regions also recorded from the eemian. mysella tumidula (jeffreys 1867) distribution. this species is recorded by jensen & knudsen (1995) as being part of the danish fauna, although it seems to have a purely lusitanian distribution (poppe & goto 1993). the species has no subfossil occurrence. montacuta substriata (montagu 1803) distribution. w iceland, the faeroes, norway north of lofoten, and south to the mediterranean. in danish waters rarely found in the north sea, skagerrak and kattegat. occurrence. the boreal and lusitanian regions. habitat. off iceland the vertical range is between 31 and 165 m (madsen 1949). in danish waters out to depths of around 700 m (skagerrak) reported as a commensal on spatangus purpureus (jensen & spärck 1934) and from the north sea also on echinocardium flavesens at depths from 30 to 100 m (petersen 1977). subfossil finds. none. fig. 64. mysella bidentata (montagu 1803). skagen 4, 27.0–27.5 m b.s., lab. no. 352,93. × 9.6. to the left a specimen seen from the right, and to the right a left valve. mguh 25374. geus bulletin no 3.pmd 28-06-2004, 08:4575 76 tellimya ferruginosa (montagu 1803) fig. 65 distribution. around iceland, the faeroes, norway north of lofoten, and south to the mediterranean. in danish waters the species extends from the north sea, limfjord, and skagerrak into the kattegat and øresund (jensen & spärck 1934). occurrence. the subarctic, boreal and lusitanian regions. habitat. in general the species is most common just below the tidal zone, which according to poppe & goto (1993) is the preferred habitat of echinocardium with which t. ferruginosa is often associated. however, the species is also found on brissopsis lyrifera or living by itself (jensen & spärck 1934). accordingly, the depth range may vary, around iceland being between 32 and 80 m (madsen 1949). subfossil finds. the limfjord, north sea, vendsyssel and skagen regions, holocene. the records from the skagen well are from the subboreal and subatlantic. furthermore, the species has been recorded from the bælt sea and north sea during the eemian. mysella dawsoni (jeffreys 1864) distribution. petersen (1888, p. 154) mentioned a single find from the limfjord, and the species is mentioned by petersen (1986a) on the basis of the tables on molluscan finds in the limfjord from danmarks fiskeriog havundersøgelser (petersen 1976). however, it is not cited among the recent danish species by jensen & knudsen (1995). the species is mentioned from w greenland (thorson 1951) and south to the mediterranean (poppe & goto 1993), and also from spitsbergen and the west coast of norway (jensen & spärck 1934). occurrence. the arctic, subarctic, boreal and lusitanian regions. habitat. deep-living species, but mentioned from a depth of 5 m by posselt & jensen (1898) and in sandy bottom in west greenland. subfossil finds. none. tellimya tenella (lovén 1846) cited only from the kattegat, including the øresund, by jensen & knudsen (1995). no subfossil finds. will not be considered further. turtonia minuta (fabricius 1780) fig. 66 distribution. w greenland, around iceland, the faeroes, norway north of lofoten, and south to the mediterranean. occurrence. the subarctic, boreal and lusitanian regions. habitat. lives in the tidal zone among plants and algae on rocks (poppe & goto 1993). off iceland in the tidal zone all around the island, but also down to a depth of 50 m (petersen 1968). only subfossil finds. the limfjord, vendsyssel and fig. 65. tellimya ferruginosa (montagu 1803). skagen 4, 27.0– 27.5 m b.s., lab. no. 352,93. × 9.6. to the left interior of a right valve, and to the right a specimen seen from the left. mguh 25375. fig. 66. turtonia minuta (fabricius 1780). skagen 3, 39.85– 40.02 m b.s., lab. no. 711,93. × 40. right valve. mguh 25376. geus bulletin no 3.pmd 28-06-2004, 08:4576 77 skagen regions, holocene. recorded from the subatlantic in the skagen well. lepton nitidum (turton 1822) distribution. from the faeroes, western norway, and south to the mediterranean (jensen & spärck 1934; madsen 1949). only recently recorded from the northern kattegat and øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from 18 to 216 m, commensal on the crustaceans upogebia deltaura and gebia stellata (nordsieck 1969, p. 89). subfossil finds. the limfjord, north sea and vendsyssel regions, holocene. recorded from the eemian in the bælt sea and north sea areas. lepton squamosum (montagu 1803) distribution. from the west coast of norway and south to spain. only shells have been found in danish waters (northern kattegat) mentioned by jensen & knudsen (1995). occurrence. the boreal and lusitanian regions. habitat. from 10 to 120 m deep it lives in and around the burrows of the crustaceans (upogebia deltaura and u. stellata) on mud or gravel bottoms (poppe & goto 1993). subfossil finds. none. devonia perrieri (malard 1904) distribution. from the british isles south to spain. in danish waters recorded from the northern kattegat (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. this species is a commensal on leptosynapta inhaerens (nordsieck 1969, p. 95). subfossil finds. none. kellia suborbicularis (montagu 1803) distribution. s and w iceland, the faeroes, norway off the lofoten islands and south to the mediterranean. in danish waters recorded from the northern kattegat and øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. intertidal to 120 m deep in crevices, shells or in borings made by other species (poppe & goto 1993). subfossil finds. recorded from the vendsyssel region, holocene. potidoma dorkiae (clark 1852) this species has been recorded only in a single find by jensen & knudsen (1995), and there are no subfossil finds from the literature, so it will not be considered further. astarte sulcata (da costa 1778) distribution. se and w greenland, s and w iceland, the faeroes, norway north of lofoten, and south to the mediterranean. the species extends from the north sea into the kattegat, but is not common (jensen & spärck 1934). occurrence. the subarctic, boreal and lusitanian regions. habitat. sublittorally from 20 m and deeper on sand, mud and gravel bottoms. subfossil finds. none. tridonta borealis schumacher 1817 distribution. this species is found in all regions of the north atlantic except the british isles (madsen 1949). however, according to petersen (1968) the species has not been recorded from the faeroes. in the danish waters the species extends from the deeper part of the northern north sea (single finds) into the kattegat, including the øresund and the bælt sea, becoming very common in the baltic (jensen & spärck 1934). occurrence. the arctic, subarctic and boreal regions. geus bulletin no 3.pmd 28-06-2004, 08:4577 78 habitat. from below the tidal zone down to 250 m on mud, sand and gravel bottoms. the species is “a common member of all the zones of the arctic macoma community, and the gomphina fluctuosa community” (ockelmann 1958, p. 83). subfossil finds. the bælt sea and limfjord regions, holocene. recorded from the vendsyssel region both in the early/middle weichselian (not in the skærumhede sequence) and the late weichselian (the older yoldia sea and the younger yoldia sea respectively). tridonta elliptica (brown 1827) distribution. this species is found in all regions of the north atlantic, including the british isles (madsen 1949). petersen (1968) specified the occurrence of this species to be in the clyde sea and off the isle of man only. in the danish waters it occurs in the kattegat, øresund, the bælt sea and the baltic, where it becomes as common as t. borealis (jensen & spärck 1934). occurrence. the arctic, subarctic and boreal regions. considering the occurrences on the west coast of britain it is also in the northern part of the lusitanian region, but here probably mostly in deeper water. habitat. the vertical range for this species is off iceland 6 to 300 m (madsen 1949), and off the east greenland coast it is most often taken between about 5 and 50 m, being abundant locally within the arctic macoma community (ockelmann 1958). subfossil finds. recorded from the north sea during the eemian, and in vendsyssel during the early/middle weichselian, being part of the older yoldia sea found in the skærumhede sequence (jessen et al. 1910). tridonta montagui (dillwyn 1817) distribution. e and w greenland, around iceland, spitsbergen, the faeroes, norway form north of lofoten, and south to the clyde sea and isle of man (petersen 1968), also recorded from the bay of biscay. it is present in the north western part of the north sea (petersen 1977) and common in the kattegat, øresund and extending into the bælt sea (jensen & spärck 1934). occurrence. the arctic, subarctic, boreal and lusitanian regions. habitat. the vertical range off iceland is 7–150 m (madsen 1949) and in the north sea sampled at depths between 40 and 75 m on mixed bottom in the trenches around the dogger bank (petersen 1977, p. 226). subfossil finds. recorded from the north sea region during the eemian and from the vendsyssel area during the early/middle weichselian (the older yoldia sea). acanthocardia echinata (linnaeus 1758) fig. 67 distribution. s and w iceland, norway north of lofoten, and south to the mediterranean (madsen 1949). in danish waters the species extends from the north sea (petersen 1977) into the skagerrak, limfjord, and kattegat regions and the øresund (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. from 4 to 350 m deep on mud, sand and gravel bottoms (poppe & goto 1993). subfossil finds. the kattegat, limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen well recorded from the subboreal and subatlantic. there are eemian records from the bælt sea, kattegat, north sea and vendsyssel regions. parvicardium exiguum (gmelin 1791) distribution. norway north of lofoten and south to the mediterranean (madsen 1949). in danish waters common in bays and fjords, including the limfjord. considering all the variations belonging to the same species it extends into the bælt sea (petersen 1888). occurrence. the boreal and lusitanian regions. fig. 67. acanthocardia echinata (linnaeus 1758). skagen 4, 27.0–27.5 m b.s., lab. no. 352,93. × 20. right valve. mguh 25377. geus bulletin no 3.pmd 28-06-2004, 08:4578 79 habitat. in general occurring from low tide to about 55 m deep (poppe & goto 1993); however, according to rasmussen (1973) the species has its main occurrence along the shores and is associated with vegetation. subfossil finds. the bælt sea, baltic, kattegat, limfjord, north sea and vendsyssel regions, holocene. recorded also from the bælt sea, kattegat, and north sea regions during the eemian. parvicardium hauniense (petersen & russell 1971) distribution. this newly established species has been recorded from recent danish waters, but no subfossil records are at hand. parvicardium ovale (sowerby 1840) distribution. around iceland, norway north of lofoten, and south to the mediterranean (madsen 1949). in danish waters the species is found in all the regions except the baltic extending only to darss (jensen & spärck 1934). occurrence. the subarctic, boreal and lusitanian regions. habitat. according to the icelandic records (madsen 1949) found between 5 to 350 m on bottoms such as mud, sand, clay and shell gravel with stones. subfossil finds. the bælt sea, limfjord, north sea and vendsyssel regions, holocene. from the north sea and vendsyssel regions also recorded from the eemian. parvicardium scabrum (philippi 1844) distribution. from norway north of the lofoten islands, and south to the mediterranean (madsen 1949). in danish waters common in the limfjord (jensen & spärck 1934) and recorded from the kattegat, but questioned, as there may be two separate species (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone to several hundred metres deep on sand, mud and gravel bottoms (poppe & goto 1993). subfossil finds. the bælt sea, kattegat, limfjord, north sea and vendsyssel regions, holocene. recorded from the north sea during the eemian. plagiocardium papillosum (poli 1795) distribution. from the english channel south into the mediterranean (poppe & goto 1993). occurrence. the lusitanian region. habitat. from 1 to 60 m deep on rough sand and gravel bottoms. only subfossil finds. recorded from the eemian in the north sea region. parvicardium minimum (philippi 1836) fig. 68 distribution. s and w iceland, the faeroes, norway north of lofoten, and south to the mediterranean at greater depths. in danish waters common in the deeper part of the skagerrak extending into the kattegat, including the øresund (petersen 1888). the occurrence in the limfjord is questioned (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from 4 to 161 m on mud, sand and gravel bottoms (poppe & goto 1993). however, according to madsen (1949) the species has around iceland only been found at depths of more than 75 m, and the bottoms are recorded as sand with shells and stones or as ooze and clay. in danish waters the species prefers fig. 68. parvicardium minimum (philippi 1836). geus collection. north of skagen, denmark. × 9.6. specimen seen from the right. mguh 25378. geus bulletin no 3.pmd 28-06-2004, 08:4579 80 depths of more than about 30 m, and it occurs at the greatest depth (petersen 1888). subfossil finds. the north sea and skagen regions, holocene. recorded from the skagen well during the preboreal–boreal, the atlantic, subboreal and subatlantic. from the vendsyssel region recorded from the eemian. cerastoderma edule (linnaeus 1758) distribution. norway north of the lofoten islands, and south to the mediterranean (madsen 1949). in danish waters found in all regions (petersen 1888, p. 136 – who already stressed that it is a very variable species). occurrence. the boreal and lusitanian regions. habitat. this is a shallow-water infaunal species – intertidal to few metres deep, but in the baltic occurring also at 20–30 m depths (jensen & spärck 1934) on sandy bottoms. subfossil finds. the bælt sea, baltic, kattegat, limfjord, north sea and vendsyssel regions, holocene. from the eemian recorded in the bælt sea, baltic, kattegat and north sea regions. cerastoderma glaucum (poiret 1789) distribution. from the west coast of norway south to the mediterranean (poppe & goto 1993, pp. 95–96). in danish waters the species extends into the baltic (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. this is a shallow-water species on sand and mud bottoms. however, according to studies on subfossil material (rasmussen 1973, p. 298–302), the associated fauna indicates a tidal estuarine environment for the danish material of an ertebølle age in the isefjord. subfossil finds. the baltic, kattegat and limfjord regions, holocene. however, the species identifications on the subfossil material recorded through time should be taken with some reservation on the basis of the great difficulties connected with recent species identifications. clinocardium ciliatum (fabricius 1780) distribution. w and e greenland, spitsbergen, around iceland and norway north of the lofoten islands. from the faeroes only at depths exceeding 400 m (petersen 1968). occurrence. the arctic, subarctic and boreal (highboreal) regions. habitat. from the tidal zone down to 700 m, off iceland occurring on ooze, mud, clay, sand and mixed bottoms (madsen 1949). mainly found in the arctic macoma community (ockelmann 1958). only subfossil finds. from the vendsyssel region recorded both from the eemian and from the early/middle weichselian (the older yoldia sea). laevicardium crassum (gmelin 1791) distribution. from norway north of the lofoten islands and south to the mediterranean. in danish waters only recorded from the north sea (petersen 1977) and the northern part of the kattegat (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone down to 183 m deep on sand, mud or gravel bottoms. subfossil finds. the vendsyssel region, holocene. from the eemian recorded in the kattegat region. serripes groenlandicus (bruguière 1798) distribution. w and e greenland, spitsbergen, around iceland and norway north of lofoten. from the faeroes only at depths exceeding 400 m (petersen 1968). occurrence. the arctic, subarctic and boreal (highboreal) regions. habitat. from 0 to 1 m to 120 m deep on clay and mud, but also sand and gravel are recorded (madsen 1949). subfossil finds. from the vendsyssel area during the eemian and early/middle weichselian. geus bulletin no 3.pmd 28-06-2004, 08:4580 81 mactra stultorum (linnaeus 1758) fig. 69 distribution. from the british isles and denmark south to the mediterranean (petersen 1968). in the danish waters the species is common in the north sea and skagerrak, extending into the limfjord (petersen 1986a), kattegat and øresund, although here only juveniles are present (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone down to 60 m in clean sand. subfossil finds. the limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen well recorded from the subatlantic. during the eemian recorded from the bælt sea region. lutraria lutraria (linnaeus 1758) distribution. norway off the lofoten islands and south to the mediterranean. since 1990 live specimens have been taken in danish waters near frederikshavn and at the skagerrak coast (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. intertidal down to 100 m, lives at depths up to 35 cm, burrowing in sand, sand–mud or gravel bottoms (poppe & goto 1993). subfossil finds. the limfjord and vendsyssel regions, holocene. spisula elliptica (brown 1827) distribution. around iceland, the faeroes, norway north of lofoten, and south to the english channel and gibraltar. the species occurs in the north sea, extending into the limfjord and kattegat and the øresund (jensen & knudsen 1995). occurrence. the subarctic, boreal and lusitanian regions. habitat. between 20 and 200 m deep in mud, sand and gravel bottoms. subfossil finds. the limfjord, north sea and vendsyssel regions, holocene. from the bælt sea region recorded during the eemian. spisula solida (linnaeus 1758) distribution. s and w iceland, and south to the mediterranean (petersen 1968). in danish waters very common offshore from the west coast of jylland (jensen & knudsen 1995 – as recorded by petersen 1977, fig. 25). the species extends into the limfjord (petersen 1986a) and kattegat, including the øresund and the bælt sea regions (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone down to 100 m on sandy bottoms. in the north sea found at 15–40 m depths in coarse sand (petersen 1977). subfossil finds. the limfjord, north sea and vendsyssel regions, holocene. recorded from the kattegat during the eemian. spisula subtruncata (da costa 1778) fig. 70a, b distribution. norway north of the lofoten islands, and south to the mediterranean (petersen 1968). in danish waters the species extends from the north sea into the kattegat, where it is also common. the species has been recorded both from the øresund and the bælt sea (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone down to 200 m deep fig. 69. mactra stultorum (linnaeus 1758). skagen 4, 15.0–15.5 m b.s., lab. no. 340,93. × 4.8. hinges of two right valves. mguh 25379. geus bulletin no 3.pmd 28-06-2004, 08:4581 82 in mud and sand (poppe & goto 1993). in danish waters common at depths out to 20–30 m in sand (jensen & spärck 1934). subfossil finds. the bælt sea, limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen well recorded from the atlantic, subboreal and subatlantic; in the subatlantic occurring in huge quantities. from the eemian recorded from the bælt sea, baltic and kattegat regions. solecurtus chamasolen (da costa 1778) distribution. norway off the lofoten islands, and south to the mediterranean. in the danish waters shells have been found near frederikshavn (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from 5 to 400 m deep on muddy bottoms. subfossil finds. none. solecurtus scopula (turton 1822) distribution. from the british isles and south to the mediterranean. shells recorded from the northern kattegat (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. mainly offshore to 110 m deep in muddy sand and on clean gravel bottoms (poppe & goto 1993). subfossil finds. none. ensis americanus gould 1870 distribution. at present occurring down to the øresund. however, the species has recently accidentally been transported to western europe from the east coast of north america (jensen & knudsen 1995), so it will not be further considered. ensis arcuatus (jeffreys 1865) distribution. the faeroes, the british isles and south to spain. in danish waters shells have been recorded from the kattegat region, including the øresund. occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone to 40 m deep in sand and gravel bottoms. subfossil finds. none. ensis ensis (linnaeus 1758) distribution. the faeroes, norway north of lofoten, and south to the mediterranean (petersen 1968). the species extends from the north sea east to the limfjord and kattegat regions, including the øresund (jensen & spärck 1934; jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone to about 80 m deep, burrowing in fine sand (poppe & goto 1993). from the north sea mainly recorded from depths of 20–30 m on the dogger bank and along the west coast of jylland (petersen 1977, fig. 40). fig. 70. a: spisula subtruncata (da costa 1778). skagen 4, 15.0–15.5 m b.s., lab. no. 340,93. × 4.8. right valve interior. mguh 25380. b: spisula subtruncata (da costa 1778). skagen 4, 15.0–15.5 m b.s., lab. no. 340,93. × 20. serrated surfaces of paired right valve laterals. mguh 25380. geus bulletin no 3.pmd 28-06-2004, 08:4582 83 subfossil finds. the limfjord, north sea and vendsyssel areas, holocene. during the eemian recorded from the bælt sea, north sea and the vendsyssel regions. ensis siliqua (linnaeus 1758) distribution. norway from north of lofoten, and south to the mediterranean. in danish waters, the species is recorded from the north sea, kattegat and øresund (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone to about 70 m deep in fine sand. subfossil finds. none. phaxas pellucidus (pennant 1777) fig. 71 distribution. norway off the lofoten islands, and south to the mediterranean (madsen 1949). in danish waters the species extends into the limfjord and kattegat, øresund and bælt sea regions (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. found offshore between 4 and 150 m deep in sand, mud and gravel bottoms (poppe & goto 1993). in the north sea it is abundant in the whole area, mainly from depths of 30–50 m (petersen 1977). subfossil finds. the limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen well recorded from the subboreal and subatlantic. during the eemian found in the north sea and vendsyssel regions. angulus tenuis (da costa 1778) fig. 72 distribution. norway off the lofoten islands, and south to the mediterranean (madsen 1949). in danish waters the species extends from the north sea, where it is common in shallow waters, into the kattegat and limfjord regions (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. common in shallow water down to 10–20 m deep in fine sand. subfossil finds. the bælt sea, limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen well recorded from the subatlantic. during the eemian found in the bælt sea and north sea regions. arcopagia crassa (pennant 1778) distribution. from norway south to west africa. only one record from danish waters other than empty shells (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone down to 150 m deep in sand, mud and shell gravel bottoms (poppe & goto 1993). subfossil finds. none. fig. 71. phaxas pellucidus (pennant 1777). skagen 4, 27.0–27.5 m b.s., lab. no. 352,93. × 9.6. view of the inside of a left valve above and a right valve below (valves not paired). mguh 25381. fig. 72. angulus tenuis (da costa 1778). skagen 4, 6.0–6.5 m b.s., lab. no. 331,93. × 4.8. view of the inside of a right valve. mguh 25382. geus bulletin no 3.pmd 28-06-2004, 08:4583 84 tellina donacina linnaeus 1758 distribution. from the shetlands over the british isles and south into the mediterranean (poppe & goto 1993). the species occurs in the southern north sea and has been recorded off edinburgh (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. from the low-tide mark to 200 m deep in sand, mud and gravel bottoms. only subfossil finds. recorded from the north sea during the eemian. fabulina fabula (gmelin 1791) fig. 73a, b distribution. norway of the lofoten islands, and south to the mediterranean (madsen 1949). the species extends into danish waters from the north sea to the kattegat and limfjord regions (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone to 50 m deep in fine sand. in danish waters common on sand between 5– 25 m, but it has been taken at depths of 40 m (jensen & spärck 1934). subfossil finds. the limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen well recorded from the subatlantic. from the eemian recorded from the north sea. tellina pygmaea (lovén 1846) fig. 74 distribution. the faeroes, norway north of lofoten, and south to the mediterranean (madsen 1949). the species extends from the north sea into the kattegat and øresund (jensen & knudsen 1995). uncertain in the records from the limfjord (collin 1884, p. 113). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone to depths of 100 m. in the north sea found at depths of 30–50 m on hard bottoms (petersen 1977). subfossil finds. from the skagen well recorded from the subatlantic. gastrana fragilis (linnaeus 1758) distribution. from the british isles including the shetlands, and south to the mediterranean. fig. 73. a: fabulina fabula (gmelin 1791). skagen 4, 6.0–6.5 m b.s., lab. no. 331,93. × 4.8. view of the inside of a right valve. mguh 25383. b: × 4.8. right valve exterior with diagonal lines running from the upper right to the lower left superimposed upon sculpture of concentric lines. mguh 25383. fig. 74. tellina pygmaea (lovén 1846). skagen 4, 27.0–27.5 m b.s., lab. no. 352,93. × 40. right valve. mguh 25384. geus bulletin no 3.pmd 28-06-2004, 08:4584 85 occurrence. mainly the lusitanian region. habitat. from below the tidal zone to a depth of 750 m in sand, mud and gravel bottoms (poppe & goto 1993). subfossil finds. recorded from the bælt sea, kattegat and the north sea during the eemian. macoma balthica (linnaeus 1758) distribution. w greenland, norway from north of lofoten, and south to the british isles (petersen 1968) and spain (poppe & goto 1993). the species is found in all the regions and extends far into the baltic, but it is not common in the north sea region from blåvandshuk and north to skagen (jensen & spärck 1934). occurrence. the subarctic, boreal and lusitanian regions. habitat. this is a shallow-water species, but in the baltic occurs also at depths of more than 50 m on soft bottoms (muus 1967, p. 163). the species is also the characterising animal of the petersen macoma balthica community so often found in the inner danish waters. however, muus (op. cit., pp. 215–217) discussed the problem concerning this community in further detail and concluded that petersen’s community can be considered a serviceable way of giving a brief description of a faunal region for other marine biologists. subfossil finds. the bælt sea, baltic, kattegat, limfjord, north sea and vendsyssel regions, holocene. recorded from the bælt sea, baltic, and the north sea during the eemian, and from the vendsyssel area during the late weichselian. macoma calcarea (gmelin 1791) distribution. w and e greenland, spitsbergen, around iceland, the faeroes, norway from north of lofoten (madsen 1949), and south into the north sea, kattegat, bælt sea and baltic (jensen & spärck 1934). occurrence. the arctic, subarctic and boreal regions. habitat. from the intertidal zone down to several hundred metres in the southern part of the distribution area. the species is the characteristic animal of the arctic macoma community (thorson 1957). subfossil finds. from the vendsyssel region, holocene. recorded from the baltic and the vendsyssel regions during the eemian, from the kattegat and vendsyssel regions during the early/middle weichselian (the older yoldia sea stage), and finally from the vendsyssel region also during the late weichselian (the younger yoldia sea stage). macoma torelli (steenstrup) jensen 1904 distribution. e and w greenland and spitsbergen. according to ockelmann (1958) this species is regarded as having a high-arctic main distribution and being restricted to the n atlantic sector. occurrence. the arctic region. habitat. recorded rarely from greenland sublittorally out to 90 m deep on clay and gravel (ockelmann 1958). subfossil finds. recorded from the vendsyssel region during the late weichselian (the younger yoldia sea). macoma loveni (steenstrup) jensen 1904 distribution. w and e greenland and spitsbergen. according to ockelmann (1958), the main distribution is high-arctic with low-arctic outposts. occurrence. the arctic and subarctic regions. habitat. at east greenland the vertical range of the species has been recorded from 4–5 m to 207 m, and the species has been taken from various types of bottoms: clay, mud, sand, gravel and stony ones (ockelmann 1958). subfossil finds. the vendsyssel region during the late weichselian (the younger yoldia sea). donax vittatus (da costa 1778) fig. 75 distribution. norway off the lofoten islands, and south to the mediterranean (madsen 1949). this species is found on the southern part of the dogger bank and along the west coast of jylland (petersen 1977) but not in the inner danish waters (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. geus bulletin no 3.pmd 28-06-2004, 08:4585 86 habitat. from the tidal zone down to 20 m in clean sand. one of the few molluscan species well suited to live in the coastal zone of sandy beaches. subfossil finds. the limfjord on old beach ridges facing the skagerrak (petersen 1976), north sea, vendsyssel and skagen regions, holocene. from the skagen well recorded from the subatlantic. found in the north sea during the eemian. gari depressa (pennant 1777) distribution. norway off the lofoten islands, and south to the mediterranean. occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone down to about 50 m in sand, mud and gravel bottoms. subfossil finds. the vendsyssel region, holocene. gari fervensis (gmelin 1791) fig. 76 distribution. w greenland, the faeroes, norway north of lofoten, and south to the mediterranean (petersen 1968). in the north sea the species is most common on the dogger bank (petersen 1977) and extends through the skagerrak into the kattegat and øresund (jensen & spärck 1934). occurrence. the subarctic, boreal and lusitanian regions. habitat. from the tidal zone to a depth of 110 m in coarse sand and shell gravel bottoms (poppe & goto 1993). in danish waters on mixed bottoms and sand at depths of 15–40 m (jensen & spärck 1934). subfossil finds. the limfjord, vendsyssel and skagen regions, holocene. recorded from the subatlantic in the skagen well. gari tellinella (lamarck 1818) distribution. s and w iceland, the faeroes, norway north of lofoten, and south to the mediterranean. the species has been taken from the northern kattegat (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone to a depth of 460 m in coarse sand, gravel and stone bottoms. subfossil finds. none. scrobicularia plana (da costa 1778) distribution. norway off the lofoten islands, and south to the mediterranean (madsen 1949). the species extends into the bælt sea at kiel and warnemünde (jensen & spärck 1934), and is also recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone to about 30 m in clay or muddy bottoms, often in estuaries (poppe & goto 1993). jensen & knudsen (1995) point to the occurrences in the wadden sea and to the sensitivity to severe winters, living in such shallow-water environments. fig. 75. donax vittatus (da costa 1778). geus collection. holland. × 4.8. view of the inside of a right valve. mguh 25385. fig. 76. gari fervensis (gmelin 1791). skagen 4, 23.0–23.5 m b.s., lab. no. 348,93. × 4.8. right valve. mguh 25386. geus bulletin no 3.pmd 28-06-2004, 08:4586 87 subfossil finds. the bælt sea, baltic, kattegat, limfjord, north sea and vendsyssel regions, holocene. during the eemian recorded from the bælt sea, baltic, kattegat, and north sea regions. abra alba (wood 1802) fig. 77 distribution. norway off the lofoten islands, and south to the mediterranean (madsen 1949). in danish waters it is found most abundantly in the shallow-water parts of the north sea (petersen 1977) and extends through the skagerrak, limfjord and kattegat and øresund into inner danish waters such as the bælt sea and the baltic to neustadt, where it is the typical bottom animal (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. from the infralittoral zone to a depth of 65 m in sand, mud or muddy gravel (poppe & goto 1993). in danish waters common at depths of 3–8 m and out to 15–20 m in soft bottoms (jensen & spärck 1934). subfossil finds. the bælt sea, kattegat, limfjord, north sea, vendsyssel and skagen regions, holocene. in the skagen well recorded from the subboreal and subatlantic. during the eemian recorded from the bælt sea, north sea and vendsyssel regions. abra nitida (müller 1776) fig. 78 distribution. recorded from s and w iceland, the faeroes, norway north of lofoten, and south to the mediterranean (madsen 1949). in danish waters recorded from the north sea, skagerrak, limfjord and kattegat (jensen & spärck 1934), including the øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. mainly offshore to depths of 200 m in sandy mud, mud or gravel bottoms (poppe & goto 1993). however, in danish waters such as the north sea on soft to mixed bottoms at depths of 40–70 m (petersen 1977) and in the skagerrak at depths of 100–300 m extending into the kattegat, including the øresund, on soft bottoms (jensen & spärck 1934). subfossil finds. the limfjord, north sea, vendsyssel and skagen regions, holocene. in the skagen well recorded from the subboreal and subatlantic. during the eemian recorded from the vendsyssel region. abra prismatica (montagu 1803) fig. 79 distribution. s and w iceland, the faeroes, norway north of lofoten, and south to the mediterranean (madsen 1949). in the danish waters the species extends from the north sea and skagerrak into the kattegat and øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from the infralittoral zone to 400 m deep in sand or muddy sand (poppe & goto 1993). in the north sea most abundant at depths deeper than 50 m on mixed bottoms (petersen 1977). subfossil finds. the limfjord, vendsyssel and skagen regions, holocene. from the skagen well recorded from the subboreal and subatlantic. during the eemian refig. 77. abra alba (wood 1802). geus collection. storebælt, denmark. × 4.8. specimen seen from the left. mguh 25387. fig. 78. abra nitida (müller 1776). geus collection. hellebæk, denmark. × 4.8. specimen seen from the left. mguh 25388. geus bulletin no 3.pmd 28-06-2004, 08:4587 88 corded from the bælt sea, kattegat, north sea, and vendsyssel regions. abra segmentum (récluz 1843) distribution. from the west coast of france into the mediterranean. occurrence. the lusitanian region. habitat. in the infralittoral zone in sandy mud (poppe & goto 1993). it seems to be connected with the shallow-water environment, also with brackish water (jensen & spärck 1934). only subfossil finds. recorded from the bælt sea and the north sea during the eemian. arctica islandica (linnaeus 1767) fig. 80 distribution. around iceland, the faeroes, norway north of lofoten, and south to the bay of biscay. in the danish waters, including the limfjord, the species extends from the north sea and skagerrak as far as the baltic (to bornholm) (jensen & spärck 1934). occurrence. the subarctic, boreal and (lusitanian) regions. however, the species tends to live more deeply in the southern part of its range (poppe & goto 1993). habitat. intertidal to 482 m in mud, sand or gravel bottoms. in the north sea mainly from depths deeper than 40 m and from mixed bottoms (petersen 1977). in inner danish waters often at depths from 10–15 to 50–60 on clay or clayey bottoms (jensen & spärck 1934). according to badarsson (1920), fishermen say that the species occurs in very shallow water, just below the low-water mark, in winter living deeply burrowed in the substrate, but in the summer often lying in abundance on the bottom. subfossil finds. the kattegat, limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen well recorded in the subatlantic. in the bælt sea, baltic and north sea (there are (single find) records from the eemian and the vendsyssel region during the late weichselian (the younger yoldia sea). kelliella miliaris (philippi 1844) fig. 81 distribution. from norway off the lofoten islands, and south to the mediterranean. from danish waters recorded from the skagerrak. according to jensen & spärck (1934), very common in the deeper part of the skagerrak. jensen & knudsen (1995) mentioned a single finding from the southern kattegat. fig. 79. abra prismatica (montagu 1803). geus collection. iceland. × 4.8. specimen seen from the left. mguh 25389. fig. 80. arctica islandica (linnaeus 1767). geus collection. læsø rende, denmark. × 4.8. juvenile specimen, beaks directed forwards. mguh 25390. fig. 81. kelliella miliaris (philippi 1844). skagen 3, 185.04–185.06 m b.s., core sample-102. × 20. two specimens seen from the left. inside the corroded specimen to the left, pyrite is seen. mguh 25391. geus bulletin no 3.pmd 28-06-2004, 08:4588 89 occurrence. the boreal and lusitanian regions. habitat. from 134 to 700 m deep (nordsieck 1969). subfossil finds. recorded from the eemian in the vendsyssel and skagen regions. glossus humanus (linnaeus 1758) distribution. s and w iceland, norway off the lofoten islands, and south to the mediterranean. in danish waters recorded from the north sea and kattegat, where shells are common (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. offshore beyond 7 m on bottoms of sand, sandy mud or soft mud (tebble 1966). subfossil finds. none. chamelea striatula (da costa 1778) fig. 82 distribution. the faeroes, norway north of lofoten, and south to the mediterranean (madsen 1949). in danish waters the species extends from the north sea, limfjord and skagerrak into the kattegat and øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from the infralittoral zone to 55 m deep on sand and mud bottoms (poppe & goto 1993). in the north sea mainly from 20 to 40 m on sand to mixed bottoms (petersen 1977). in the skagerrak at depths of more than 100 m, but in the kattegat at depths of less than 50 m, since the sand bottom is not to be found at deeper levels (jensen & spärck 1934). subfossil finds. the limfjord, north sea, vendsyssel and skagen areas, holocene. from the skagen well recorded from the subboreal and subatlantic. during the eemian recorded from the bælt sea, north sea and vendsyssel regions. clausinella fasciata (da costa 1778) distribution. the faeroes, norway off the lofoten islands, and south to the mediterranean. from the north sea the species extends into the kattegat (petersen 1888). occurrence. the boreal and lusitanian regions. habitat. from 4 to 110 m deep in sand, mud and gravel bottoms (poppe & goto 1993). in the north sea in hard sand (petersen 1977), and in the kattegat in gravel and sand between 15 and 30 m (petersen 1888, p. 143). subfossil finds. the limfjord, north sea and vendsyssel areas, holocene. paphia aurea (gmelin 1791) distribution. norway off the lofoten islands, and south to the mediterranean. the finds closest to danish waters are from southern norway (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone down to 36 m deep in sand, mud and gravel bottoms. only subfossil finds. the bælt sea, kattegat, limfjord, north sea and vendsyssel regions, holocene. recorded from the kattegat region during the eemian. paphia aurea senescens (cocconi 1873) distribution. the only one of our quaternary molluscs which does not live at present. also found in quaternary deposits in italy (jensen & spärck 1934). regarded as a lusitanian species according to nordmann (1913). however, poppe & goto (1993) regard the fossil valves found, for example along the coast of the netherlands and belgium, as a subspecies in which the differences fig. 82. chamelea striatula (da costa 1778). skagen 4, 21.0– 21.5 m b.s., lab. no. 346,93. × 4.8. left valve. mguh 25392. geus bulletin no 3.pmd 28-06-2004, 08:4589 90 from extant ones are minimal, and they propose that the relationship between the fossil and recent shells be restudied. from nordmann (1913) and cerulli-irelli (1908) it seems right that tapes senescens doederl. and tapes aureus var. eemiensis are identical. but as the relationship between t. senescens and t. aureus sensu stricto at the time of nordmann (1913) was not clear, the position as a not extant subspecies given by poppe and goto is followed here. only subfossil finds. the bælt sea, kattegat, and north sea regions during the eemian. tapes decussatus (linnaeus 1758) distribution. norway off the lofoten islands, and south to the mediterranean. closest to danish waters the species occurs off western norway (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone to a depth of few metres in sand or muddy-gravel bottoms. only subfossil finds. the kattegat, limfjord, north sea and vendsyssel regions, holocene. recorded from the north sea region during the eemian. timoclea ovata (pennant 1777) fig. 83a, b distribution. s and w iceland, the faeroes, norway north of lofoten, and south to the mediterranean (madsen 1949). in danish waters the species extends from the north sea into the kattegat and øresund (jensen & knudsen 1995), but has not been recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. at depths between 4 and 200 m on all types of bottoms (poppe & goto 1993). in the north sea usually deeper than 50 m and on soft bottoms (petersen 1977). subfossil finds. the limfjord, north sea, vendsyssel and skagen regions, holocene. in the skagen well recorded from the subatlantic. during the eemian found in the bælt sea, kattegat and north sea regions. venerupis rhomboides (pennant 1777) distribution. the faeroes, norway off the lofoten islands (madsen (1949) does not mention any norwegian occurrence), and south to the mediterranean (petersen 1968). one pair of united valves recorded from the northern kattegat (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone to 180 m deep (poppe & goto 1993) in gravel and mud bottoms. subfossil finds. the limfjord and vendsyssel regions, holocene. venerupis pullastra (montagu 1803) distribution. norway from north of the lofoten islands, and south to the mediterranean (madsen 1949). from danish waters recorded from the limfjord and kattefig. 83. a: timoclea ovata (pennant 1777). skagen 4, 27.0–27.5 m b.s., lab. no. 352,93. × 40. left valve. mguh 25393. b: timoclea ovata (pennant 1777). geus collection. herthas flak, denmark. × 4.8. right valve. mguh 25394. geus bulletin no 3.pmd 28-06-2004, 08:4590 91 gat regions (jensen & spärck 1934) and the øresund (jensen & knudsen 1995). the species has a common occurrence in the isefjord (rasmussen 1973). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone to 40 m deep in hard sand and muddy gravel (poppe & goto 1993). rasmussen (1973, p. 303): considers “its present common occurrence there [in the isefjord] to be a result of the disappearance of the zostera since 1933–45, as the lack of a continuous vegetation caused a change in the bottom conditions to the benefit of venerupis pullastra. undoubtedly the species has lived in interior danish waters since the stone age, being, however, rare in recent times up to 1933–34”. subfossil finds. the bælt sea, kattegat, limfjord, north sea and vendsyssel regions, holocene. recorded from the north sea region during the eemian. dosinia exoleta (linnaeus 1758) distribution. norway off lofoten, and south to the mediterranean. the species extends from the north sea (petersen 1977) into the northern kattegat (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone to 73 m deep, burrowing deeply in sand, mud and gravel bottoms (poppe & goto 1993). in the north sea found at depths of 30– 100 m, on hard bottoms around the dogger bank (petersen 1977). subfossil finds. the north sea and vendsyssel regions, holocene. dosinia lincta (montagu 1803) distribution. s and w iceland, the faeroes, norway north of lofoten, and south to the mediterranean (petersen 1968). the species extends from the north sea into the kattegat and øresund (petersen 1888). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone to 200 m deep in pure sand and fine gravel bottoms (poppe & goto 1993). in the north sea sampled at 10–100 m on soft to mixed bottoms (petersen 1977). from the kattegat recorded on mixed bottoms between 18 and 56 m (petersen 1888). subfossil finds. the north sea and vendsyssel regions, holocene. recorded from the bælt sea and the north sea during the eemian. gouldia minima (montagu 1803) distribution. the faeroes, norway off the lofoten islands, and south to the mediterranean (petersen 1968). occurrence. the boreal and lusitanian regions. habitat. from below the tidal zone to depths of over 200 m in sand, mud and fine gravel bottoms (poppe & goto 1993). only subfossil finds. recorded from the bælt sea and the north sea during the eemian. petricola pholadiformis (lamarck 1822) distribution. this species in an immigrant introduced in to europe, probably with oysters, at the end of the last century (poppe & goto 1993). according to jensen & knudsen (1995), the species occurred in 1906 in the wadden sea; the skagerrak 1905; the kattegat 1931; and the bælt sea 1943. it will not be considered any further in this work. mysia undata (pennant 1777) distribution. the faeroes, norway north of the lofoten islands, and south to the mediterranean (petersen 1968). in danish waters the species extends from the north sea (petersen 1977) into the kattegat and øresund (jensen & knudsen 1995). it is not recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from below the tidal zone to depths of 55 m in muddy sand and gravel bottoms (poppe & goto 1993). in the north sea found at depths of 40–70 m on soft to mixed bottoms (petersen 1977). subfossil finds. the limfjord and vendsyssel regions, holocene. during the eemian recorded from the north sea region. geus bulletin no 3.pmd 28-06-2004, 08:4591 92 order myoida mya arenaria linnaeus 1758 distribution. norway from north of the lofoten islands, south to the british isles (petersen 1968). the species extends from the north sea into the limfjord and the inner danish waters, including the baltic (jensen & spärck 1934). it is a late immigrant, known from europe in the plio-pleistocene (strauch 1972, pp. 135– 137) having been transferred from north america by man, presumably the vikings, and dated back to the 13th century, i.e. well before columbus (petersen et al. 1992b). occurrence. mainly in the boreal region, but with new finds further to the south on the east coast of north america (rasmussen & heard 1994). habitat. from the tidal zone down to 6–7 m deep in sandy bottoms, the species avoids high-energy coastal environments (jensen & spärck 1934). subfossil finds. the records from the bælt sea and vendsyssel might be of recent dates; only the occurrences at jerup halfway up to the skagen spit have been dated and included in the skagen area from the subatlantic. mya truncata linnaeus 1758 distribution. w and e greenland, spitsbergen, around iceland, the faeroes, norway from north of lofoten, and south to the bay of biscay (madsen 1949; poppe & goto 1993). the species extends from the north sea into the limfjord and the inner danish waters as far as the bælt sea (kiel and warnemünde) (jensen & spärck 1934). occurrence. the arctic, subarctic, boreal and lusitanian regions. habitat. from the intertidal zone down to about 75 m deep (poppe & goto 1993). however, in danish waters often at depths between 10 and 20 m (jensen & spärck 1934). in the north sea found at depths of 37– 70 m on soft mixed bottoms (petersen 1977). in east greenland the species belongs to the arctic macoma community (ockelmann 1958). however, according to jensen (1900) the typical mya truncata is not found in thearctic (g.h.petersen,personalcommunication1998). subfossil finds. the bælt sea, kattegat, limfjord, north sea and vendsyssel regions, holocene. recorded from the eemian in the bælt sea, baltic, north sea and vendsyssel areas. from the early/middle weichselian in the kattegat and vendsyssel regions, and from the vendsyssel region also in the late weichselian. corbula gibba (olivi 1792) fig. 84 distribution. norway north of lofoten, and south to the mediterranean (petersen 1968). the species extends from the north sea into the inner danish waters, including the limfjord, as far as the westernmost part of the baltic region at møn (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. from the low intertidal zone to 250 m deep anchored by a byssus on silty sand and muddy-gravel bottoms (poppe & goto 1993). however, in danish waters rarely at depths of more than about 50 m (jensen & spärck 1934). in the north sea sampled at 35–50 m depths on mixed bottoms (petersen 1977). subfossil finds. the bælt sea, baltic, kattegat, limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen well recorded from the subboreal and subatlantic. during the eemian recorded from the bælt sea, baltic, kattegat, north sea and vendsyssel regions. fig. 84. corbula gibba (olivi 1792). skagen 3, 37.00–37.25 m b.s., lab. no. 710,93. × 9.6. to the upper left a specimen seen from the left, and to the right a view of the inside of a right valve. mguh 25395. geus bulletin no 3.pmd 28-06-2004, 08:4592 93 hiatella arctica (linnaeus 1758) fig. 85 distribution. w and e greenland, spitsbergen, around iceland, the faeroes, norway from north of lofoten, and south to the mediterranean (madsen 1949). following jensen & spärck (1934) who regard the records of hiatella as one species, it is widely spread but not always common in all the danish waters, it extends to kiel in the bælt sea. occurrence. the arctic, subarctic, boreal and lusitanian regions. habitat. “from the intertidal zone down to almost 1400 m fixed by its byssus on or in all kinds of substrate on all types of bottoms. also found in holes previously bored by other species” (poppe & goto 1993, p. 130). however, petersen (1977, p. 228) states: “both the systematics and nomenclature are insufficiently investigated for this genus”. so with regard to the subfossil finds, the questions on species are even more difficult, as seen from petersen (1986b, figs 2, 3), where forms with different habitat such as hiatella cf. byssifera are found fixed on a stone taken as a grab sample in kejser franz josephs fjord, east greenland, and as traces of hiatella arctica in the saxicava sand of late weichselian age in vendsyssel. however, here símonarson et al. (1998) is followed, relating the more widely used and less specific name hiatella arctica. subfossil finds. the bælt sea, limfjord, north sea, vendsyssel and skagen areas, holocene. from the skagen well records from the subboreal and subatlantic. from the eemian recorded from the kattegat, north sea, vendsyssel and skagen regions. in the kattegat and vendsyssel regions finds from the early/middle weichselian (the older yoldia sea) and in the vendsyssel region from the late weichselian (the younger yoldia sea). hiatella rugosa (linnaeus 1758) distribution. jensen & knudsen (1995) include this as a separate species and take it as part of the recent danish shell-bearing fauna. records from the literature on subfossil finds are therefore considered here. according to poppe & goto (1993), found from norway south to the mediterranean. ockelmann (1958, p. 135 ff.), discussing the hiatella taxonomy at some length, concludes that reservations must be made as to future separations of the hiatella forms (h. arctica incl. of h. gallicana and h. pholadis) occurring in the northern hemisphere into valid species. occurrence. the arctic, subarctic, boreal and lusitanian regions. subfossil finds. the kattegat and vendsyssel regions, holocene. saxicavella jeffreysi winckworth 1930 fig. 86 distribution. norway off the lofoten islands, and south to the mediterranean (madsen 1949). the species extends into the kattegat and øresund, although rare (jensen & knudsen 1995). it has been recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. offshore between 7 and 240 m deep in sand, mud or gravel bottoms (poppe & goto 1993). in the fig. 85. hiatella arctica (linnaeus 1758). skagen 3, 37.00–37.25 m b.s., lab. no. 710,93. × 20. left valve. mguh 25396. fig. 86. saxicavella jeffreysi winckworth 1930. skagen 4, 20.0– 20.5 m b.s., lab. no. 345,93. × 4.8. right valve. mguh 25397. geus bulletin no 3.pmd 28-06-2004, 08:4593 94 danish waters often at depths between 25–50 m (jensen & spärck 1934). subfossil finds. the limfjord, north sea, vendsyssel and skagen regions, holocene. in the skagen well recorded from the subatlantic. during the eemian recorded from the north sea region. panomya arctica (lamarck 1818) distribution. s and w iceland, norway from north of lofoten, and south to the british isles and denmark (madsen 1949). however, petersen (1968) refers to empty shells from n iceland and occurrences in the clyde sea and off the isle of man. in danish waters only once taken alive near skagen, otherwise shells only, but found as far south as øresund (jensen & knudsen 1995). occurrence. mainly boreal, but outposts into the subarctic (subfossil?) and lusitanian regions (strauch 1972). habitat. from the intertidal zone down to 300 m buried in mud or sand (poppe & goto 1993). subfossil finds. none. barnea candida (linnaeus 1758) fig. 87 distribution. norway off the lofoten islands, and south to the mediterranean (madsen 1949). in the danish waters, including the limfjord, the species extends into the bælt sea as far as kiel (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. from the low intertidal zone to about 30 m deep, the species bores in semi-hard substrates such as clay (poppe & goto 1993). subfossil finds. the bælt sea, limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen well recorded from the subatlantic. during the eemian found in the bælt sea and north sea regions. pholas dactylus linnaeus 1758 fig. 88a, b distribution. norway off the lofoten islands, and south to the mediterranean (madsen 1949). in danish waters only found to frederikshavn and the limfjord (petersen 1986a; jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone to a depth of 10 m boring in different substrates, preferring clay bottoms (poppe & goto 1993). subfossil finds. the limfjord, vendsyssel and skagen regions, holocene. from the skagen well recorded from the subatlantic. zirfaea crispata (linnaeus 1758) distribution. around iceland, norway from north of the lofoten islands, and south to the bay of biscay (madsen 1949). in the danish waters, including the limfjord, extending into the bælt sea as far as kiel (jensen & spärck 1934). occurrence. mainly boreal with outposts into the lusitanian region. habitat. from the low tide line to about 7 m deep, boring in semi-hard substrates (poppe & goto 1993). in the danish waters the species has a wide extension, depending on the bottom substrates (jensen & spärck 1934): in the north sea and skagerrak peat and chalk; in the limfjord cementstone, mo-clay, chalk, peat and clay; and the bælt sea clayey bottoms. subfossil finds. the bælt sea, limfjord, north sea and vendsyssel regions, holocene. from the bælt sea region recorded from the eemian and from the vendsyssel region during the late weichselian (the younger yoldia sea). fig. 87. barnea candida (linnaeus 1758). geus collection. off rubjerg knude, denmark. × 4.8. left valve. mguh 25398. geus bulletin no 3.pmd 28-06-2004, 08:4594 95 xylophaga dorsalis turton 1822 distribution. s and w iceland, norway from north of lofoten, and south to the mediterranean. recorded from the øresund region (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. according to madsen (1949) the occurrences at iceland have been from depths between 140–230 m found in sunken pieces of wood. subfossil finds. none. teredo navalis linnaeus 1758 distribution. norway off the lofoten islands (petersen 1968), and south to the mediterranean (madsen 1949). the species is found into the bælt sea (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. specialised wood-borers. subfossil finds. none. nototeredo norvegica (spengler 1792) distribution. s and w iceland, norway north of lofoten, and south to the mediterranean (madsen 1949). the species extends from the north sea (in driftwood) into the bælt sea as far as kiel (jensen & spärck 1934), and was recorded from the limfjord by collin (1884). occurrence. the boreal and lusitanian regions. habitat. in driftwood. subfossil finds. none. psiloteredo megotara (forbes & hanley 1848) distribution. w greenland, spitsbergen, around iceland, norway north of lofoten, and south to the mediterranean (madsen 1949). in danish waters common on the west coast of jylland and recorded from the kattegat, including øresund (jensen & knudsen 1995). occurrence. the arctic, subarctic, boreal and lusitanian regions, although rare in the last region. habitat. only in driftwood (jensen & knudsen 1995). subfossil finds. none. subclass anomalodesmata order pholadomyoida pandora glacialis leach 1819 distribution. w and e greenland and spitsbergen (madsen 1949). occurrence. the arctic and subarctic regions. habitat. from 2 to 205 m deep on mixed bottoms (ockelmann 1958). only subfossil finds. from the vendsyssel region during early/middle and late weichselian (the older yoldia sea and the younger yoldia sea respectively). fig. 88. a: pholas dactylus linnaeus 1758. skagen 4, 20.0–20.5 m b.s., lab. no. 345,93. × 9.6 fragment with umbonal reflection with septa. mguh 25399. b: pholas dactylus linnaeus 1758. geus bulletin no 3.pmd 28-06-2004, 08:4595 96 lyonsia norwegica (gmelin 1791) distribution. s and w iceland, the faeroes, norway off the lofoten islands, and south to the mediterranean. in danish waters recorded from the kattegat. occurrence. the boreal and lusitanian regions. habitat. from 20 to 250 m deep in sand, silty sand and mud bottoms. subfossil finds. recorded from the skagen region, holocene. in the skagen well recorded from the preboreal and boreal. lyonsia arenosa (möller 1842) distribution. w and e greenland, spitsbergen, and norway north of the lofoten islands. main distribution in the arctic and subarctic regions. habitat. from 3 to about 200 m on mixed bottoms (ockelmann 1958). subfossil finds. recorded from the vendsyssel region during the early/middle weichselian and the late weichselian (the older yoldia sea and the younger yoldia sea respectively). cochlodesma praetenue (pulteney 1799) fig. 89a, b distribution. s and w iceland, the faeroes, norway off the lofoten islands, and south to gibraltar (petersen 1968). the species is rare in danish waters and has only once been taken live in the kattegat (jensen & spärck 1934), although shells are found in the øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone down to 110 m in sand, mud and gravel bottoms. subfossil finds. the north sea and skagen regions, holocene. from the skagen well recorded from the subatlantic. during the eemian recorded from the skagen region. thracia convexa (wood 1815) distribution. the faeroes, norway off the lofoten islands, and south to the mediterranean (madsen 1949). in danish waters recorded from the kattegat (jensen & knudsen 1995). occurrence. the boreal and lusitanian zones. habitat. offshore down to over 800 m in mud and sand bottoms (poppe & goto 1993). in the danish waters taken between 30–80 m (jensen & spärck 1934). subfossil finds. none. thracia phaseolina (lamarck 1818) fig. 90 distribution. s and w iceland, norway off the lofoten islands, and south to the mediterranean (madsen 1949). the species extends from the north sea (petersen 1977) into the kattegat and øresund, although with few records (jensen & knudsen 1995), and it has been refig. 89. a: cochlodesma praetenue (pulteney 1799). skagen 3, 43.19–43.24, core sample-9. × 9.6. view of the inside of a fragmented right valve. mguh 25401. b: cochlodesma praetenue (pulteney 1799). geus collection. læsø rende, denmark. × 4.8. detailed view of the inside of a right valve showing the resilifer. mguh 25402. geus bulletin no 3.pmd 28-06-2004, 08:4596 97 corded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from the low intertidal zone down to 50 m in fine sand, mud or gravel bottoms (poppe & goto 1993). subfossil finds. the limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen well recorded from the subatlantic. during the eemian found in the bælt sea and north sea regions. thracia gracilis (jeffreys 1865) distribution. recorded from the atlantic (nordsieck 1969). from the danish waters found in the northern kattegat (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. subfossil finds. none. thracia villosiuscula (macgillivray 1827) distribution. s and w iceland, the faeroes, norway off the lofoten islands, and south to the british isles (petersen 1968). according to poppe & goto (1993) also found in the mediterranean. occurrence. the boreal and lusitanian regions. habitat. in the north sea sampled at depths between 20 and 50 m on hard sand (petersen 1977). subfossil finds. recorded from the north sea region during the eemian. cuspidaria cuspidata (olivi 1792) distribution. norway off the lofoten islands, and south to the mediterranean. in the danish waters found in the kattegat (jensen & spärck 1934) and øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from 20 m to 250 m deep in muddy sand and gravel bottoms (poppe & goto 1993). in danish waters at depths between 30 and 60 m (jensen & spärck 1934). subfossil finds. none. cuspidaria obesa (lovén 1846) distribution. e and w greenland, spitsbergen, around iceland, norway north of the lofoten islands, and south to the mediterranean; however, to the south only at depths greater than 400 m (madsen 1949). in danish waters recorded from the skagerrak and kattegat. occurrence. the arctic, subarctic and boreal regions. habitat. at depths from 40 to 2500 m according to madsen (1949), but in danish waters fairly common in the deeper part of the skagerrak. subfossil finds. none. fig. 90. thracia phaseolina (lamarck 1818). skagen 4, 14.0– 14.5 m b.s., lab. no. 339,93. × 9.6. fragment of left valve. mguh 25403. geus bulletin no 3.pmd 28-06-2004, 08:4597 geological survey of denmark and greenland bulletin 4, 2003, pp 45-48 45 in recent years the geological survey of denmark and greenland (geus) has, for a variety of reasons, worked intensively on implementing internet technologies. the most important aim has been to provide public and private companies with access to many of the geological databases and maps at the survey, thus substantially increasing the value of the geological data. in this way geus is implementing the intentions of the united nations aarhus-convention in respect of access to information, public participation in decision-making and access to justice in environmental matters (unec 1998). another important objective has been to improve the procedures that deal with data input, registration and quality control of the large amounts of data that geus receives from regional authorities, private drilling companies and advisers, as well as that acquired by the activities of its own geological staff. this complies with the danish government’s initiatives for making the public sector more efficient through digital data management. openness during most of its lifetime the survey has allocated substantial resources to the task of organising and securing the vast environmental data and the internet: openness and digital data management jørgen tulstrup fig. 1. coverage of danish shallow geophysical data from the gerda database. geological survey of denmark and greenland bulletin 4, 45–48 (2004) © geus, 2004 amount of geological data the survey is responsible for as a national institution. in 1926 the well data archive was established in response to a law passed by the danish parliament requiring that information from all new water supply wells should be reported to the survey. fifty years later the first version of a database for these data was built. since then the database has undergone continuous development, including two major upgrades to new technology, and is now known as the jupiter database. in the same period the quality of data has been improved considerably through close cooperation with the drilling companies who supply the data, and through quality control procedures at the survey. jupiter is today a nationwide database designed for data relating to shallow boreholes, groundwater chemistry, water levels, water resources, etc. this database contains data for more than 230 000 danish boreholes and more than 5.7 million chemical analyses of groundwater and drinking water. another nationwide database called gerda (geophysical relational database) was established in 1999 and is designed for shallow geophysical data including geoelectrical profiles, geoelectrical soundings, electromagnetic soundings, borehole logs and geophysical models. this database contains geophysical data and interpretations for more than 600 000 localities in denmark. through the years great efforts have been made into making the data as useful and accessible as possible for internal as well as external users, continually updating the databases to keep pace with the increased quantity of data and continual improvements in technology. the widespread development of the internet, and increased broad-band access, has made it possible to place data at the disposal of the general public and to make it available in an increasingly user-friendly and understandable way. 46 fig. 2. map available in a web-browser showing the position of schlumberger soundings (purple/blue circles), pulled array continuous electrical soundings (purple/blue lines), transient electromagnetic soundings (red/orange circles) and boreholes with log-data (red/green circles). the gis functionality is based on the esri product arcims. it is now possible to use a standard web browser to search for data utilising either geographic information system (gis) facilities or a more traditional alphanumeric search. with the gis interface (fig. 1) it is possible to zoom, pan and choose an object on the map, e.g. a borehole or a sounding position (fig. 2). by clicking the object the user can be guided to a website showing details of the object, for example a graphic illustration of a borehole log and a geological interpretation (fig. 3). using the alphanumeric search (fig. 4) the user will be presented with a list containing boreholes fulfilling different search criteria, e.g. name of drilling company, municipality, time of drilling, maximum hole depth, purpose of the borehole, or distance to a given point, as well as the possibility of searching on the unique borehole number. furthermore, it is now possible to search for wells that penetrate certain geological horizons. another data type available on the internet is a nationwide surface geological map (fig. 5). this map of denmark is based on field work that the survey has carried out since 1890. the map shows the geology at a depth of one metre and was originally produced at the scale of 1:25 000. digital data management to reduce the steadily increasing cost related to registration and quality control of the geological data, a digital data management system has been introduced. previously the necessary procedures included manually typing or scanning large amounts of data. however, since the data to be entered into the databases are often in a digital format when they are received, there is now the possibility of introducing complete or partly automated procedures in which computer programs handle quality control and loading of data into the databases. closer manual inspection is only necessary when data quality is considered uncertain compared to existing data. the latest internet technology also makes it possible for individual users to upload their own data. the gerda system was, from the outset, designed with the internet in mind. the data suppliers, such as geophysical companies, are able to directly upload data files. the data are quality controlled and then loaded into the database. an email message to the supplier supplies information on the progress of the procedure. after this is complete, an updated database is extracted from the central gerda database and made available to users of the data, e.g. geologists in the danish counties. uploaded data files in gerda contain information about projects, recording instruments, etc. this information is accessible to the data suppliers, enabling them to create and 47 fig. 4. web-based search form for the jupiter database. by entering search criteria the user can produce a list of boreholes. each borehole on the resulting list can be examined in detail. fig. 3. display of well-log data from the gerda database. the graphics are available as a pdf file. 48 maintain data directly on the website. the procedure ensures a very flexible and non-bureaucratic work routine. in 1999 geus merged four older databases into a single database (jupiter) to obtain a much higher level of automated treatment of the collected data. this has made it possible for geus to receive digital data reports of groundwater and drinking water quality, water resources and water level observations from the districts. the data are received in the so-called ‘standat’ format (generally used for exchange of environmental data in denmark) and is converted into the xml-format (extensible markup language). when completed, a comprehensive quality control is undertaken. the supplier automatically receives a load report by e-mail. the report is a detailed description of syntax errors, data errors, data exceeding existing limits, and information on whether substance concentrates have been registered to be diverging substantially from previous measurements. in the near future it will also be possible to accept basic borehole data digitally. geus has developed a program, pc jupiter, which will be used by the drilling companies for planning new water supply boreholes, as well as for registration of well localities, the technical data of the borehole, the observed geological stratification and test pumping. the program stores the data in a database administered locally by the drilling company. the data are also forwarded to geus where the data are validated and finally loaded into the central jupiter database. web-services the internet solutions described above are designed mainly to make data accessible on the geus website using traditional html-technology. to further increase the accessibility of the geological data, a new group of technologies known as web-services are under implementation. a web-service is a function available on a web-server, which on the request of a user or another webserver delivers data in xml-format, or maps as images (web map services). this technology will enable other authorities and companies who set up web-servers to make use of geological data without having to be concerned about keeping local copies of data complete and up-to-date. the data are retrieved on-line from geus’ systems. an example of this is a solution geus is providing for an internet portal that the danish counties are implementing. the portal provides access to a wide range of environmental data including borehole, groundwater and drinking water data. the portal will make it possible for users to view data on a map and to display graphs of time series and download data. the data will be retrieved live from the jupiter database through a web-service interface. data input to jupiter from the regional authorities will also take place through web-services. reference unec 1998: the aarhus convention. resolution on access to information, public participation in decision-making and access to justice in environmental matters, fourth ministerial conference ‘environment for europe’, aarhus, denmark, 23–25 june 1998. united nations economic commission for europe, 4 pp. author’s address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: jtu@geus.dk fig. 5. example of surface geological map of denmark available at the geus website. << /ascii85encodepages false /allowtransparency false /autopositionepsfiles true /autorotatepages /all /binding /left /calgrayprofile (dot gain 20%) /calrgbprofile (srgb iec61966-2.1) /calcmykprofile (u.s. web coated \050swop\051 v2) /srgbprofile (srgb iec61966-2.1) /cannotembedfontpolicy /warning /compatibilitylevel 1.4 /compressobjects /tags /compresspages true /convertimagestoindexed true /passthroughjpegimages true /createjdffile false /createjobticket false /defaultrenderingintent /default /detectblends true /colorconversionstrategy /leavecolorunchanged /dothumbnails false /embedallfonts true /embedjoboptions true /dscreportinglevel 0 /syntheticboldness 1.00 /emitdscwarnings false /endpage -1 /imagememory 1048576 /lockdistillerparams false /maxsubsetpct 100 /optimize true /opm 1 /parsedsccomments true /parsedsccommentsfordocinfo true 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<feffd5a5c0c1b41c0020c778c1c40020d488c9c8c7440020c5bbae300020c704d5740020ace0d574c0c1b3c4c7580020c774bbf8c9c0b97c0020c0acc6a9d558c5ec00200050004400460020bb38c11cb97c0020b9ccb4e4b824ba740020c7740020c124c815c7440020c0acc6a9d558c2edc2dcc624002e0020c7740020c124c815c7440020c0acc6a9d558c5ec0020b9ccb4e000200050004400460020bb38c11cb2940020004100630072006f0062006100740020bc0f002000520065006100640065007200200035002e00300020c774c0c1c5d0c11c0020c5f40020c2180020c788c2b5b2c8b2e4002e> /chs <feff4f7f75288fd94e9b8bbe7f6e521b5efa76840020005000440046002065876863ff0c5c065305542b66f49ad8768456fe50cf52068fa87387ff0c4ee563d09ad8625353708d2891cf30028be5002000500044004600206587686353ef4ee54f7f752800200020004100630072006f00620061007400204e0e002000520065006100640065007200200035002e00300020548c66f49ad87248672c62535f003002> /cht <feff4f7f752890194e9b8a2d5b9a5efa7acb76840020005000440046002065874ef65305542b8f039ad876845f7150cf89e367905ea6ff0c4fbf65bc63d066075217537054c18cea3002005000440046002065874ef653ef4ee54f7f75280020004100630072006f0062006100740020548c002000520065006100640065007200200035002e0030002053ca66f465b07248672c4f86958b555f3002> >> >> setdistillerparams << /hwresolution [2400 2400] /pagesize [595.000 842.000] >> setpagedevice geological survey of denmark and greenland bulletin 6, 51-71 51geological survey of denmark and greenland bulletin 5, 51–71 © geus, 2004 a new middle–upper jurassic succession on hold with hope, north-east greenland henrik vosgerau, michael larsen, stefan piasecki and jens therkelsen a succession of marine, jurassic sediments was recently discovered on hold with hope, northeast greenland. the discovery shows that the area was covered by the sea during middle–late jurassic transgressive events and thus adds to the understanding of the palaeogeography of the area. the jurassic succession on northern hold with hope is exposed in the hangingwalls of small fault blocks formed by rifting in late jurassic – early cretaceous times. it unconformably overlies lower triassic siltstones and sandstones and is overlain by lower cretaceous coarsegrained sandstones with an angular unconformity. the succession is up to 360 m thick and includes sandstones of the lower–upper callovian pelion and middle–upper oxfordian payer dal formations (vardekløft group) and heteroliths and mudstones of the upper oxfordian – lower kimmeridgian bernbjerg formation (hall bredning group). the pelion formation includes the new spath plateau member (defined herein). the palaeogeographic setting was a narrow rift-controlled embayment along the western margin of the rifted jurassic seaway between greenland and norway. it was open to marine circulation to the south as indicated by the distribution and lateral facies variations and a dominant south-westwards marine palaeocurrent direction. the pelion and payer dal formations represent upper shoreface and tidally influenced delta deposits formed by the migration of dunes in distributary channels and mouthbars over the delta front. the boundary between the two formations is unconformable and represents a late callovian – middle oxfordian hiatus. it is interpreted to have formed by subaerial erosion related to a sea-level fall combined with minor tilting of fault blocks and erosion of uplifted block crests. in late jurassic time, the sand-rich depositional systems of the pelion and payer dal formations drowned and offshore transition – lower shoreface heteroliths and offshore mudstones of the bernbjerg formation accumulated. the fault block crest forming the eastern basin margin was inundated by a rise in relative sea level. major fault activity probably occurred in latest jurassic – early cretaceous times when the major fault block originally defining the hold with hope basin was split into smaller blocks. keywords: hall bredning group, hold with hope, lithostratigraphy, north-east greenland, palaeogeography, sedimentology, shallow marine, spath plateau member, vardekløft group h.v.*, m.l., s.p. & j.t.‡, geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: mil@geus.dk present addresses: *roskilde amt, køgevej 80, dk-4000 roskilde, denmark. ‡skude & jacobsen, næstvedvej 1, dk-4760 vordingborg, denmark. geus bulletin no 5.pmd 29-10-2004, 11:1451 52 a new middle–upper jurassic succession, up to 360 m thick, was found recently on northern hold with hope, north-east greenland (fig. 1; stemmerik et al. 1997; kelly et al. 1998; larsen et al. 1998). it spans the early callovian – early kimmeridgian time interval as indicated by dinoflagellate cysts and ammonites, and consists of coarse-grained sandstones overlain by heteroliths and mudstones. the sandstone-dominated lower part of the succession assigned here to the pelion formation was originally studied by koch (1932) and maync (1949) and was tentatively given an early cretaceous age, although w. maync noted the resemblance to themiddle jurassic succession on wollaston forland. the apparent absence of jurassic sediments in the hold with hope area was explained differently by maync (1947), donovan (1957), surlyk (1977) and stemmerik et al. (1993). maync (1947) and surlyk (1977) suggested that during the jurassic the area formed a landmass between the wollaston forland basin to the north and the jameson land basin to the south, implying that the lack of sediments was primarily due to non-deposition. donovan (1957) in contrary found it 20ºw 16ºw24ºw28 76 75 76ºn 74ºn 72ºn greenland store koldewey wollaston forland milne land jameson land traill ø geographical society ø kuhn ø clavering ø fig. 2 hold with hope pdmf dcf jurassic fault dombjerg–clavering faultdcf 100 km post-devonian main faultpdmf fig. 1. map of east greenland showing the distribution of jurassic sediments and major faults. rectangle marks the investigated area on northern hold with hope, shown in more detail in figure 2. modified from koch & haller (1971) and surlyk et al. (1973). geus bulletin no 5.pmd 29-10-2004, 11:1452 53 most likely that the absence was secondary owing to pre-aptian erosion of the jurassic rocks. based on comparison with nearby clavering ø, stemmerik et al. (1993) suggested that middle–upper jurassic sediments were present in the subsurface east of the continuation of the dombjerg–clavering fault on hold with hope (fig. 1, dcf). the present investigations confirm that hold with hope formed a landmass between the wollaston forland and jameson land basins during much of the middle jurassic time interval. the discovery on hold with hope of a marine succession of early callovian – kimmeridgian age, however, shows that the land mass was flooded in late middle jurassic time and continued to be sea-covered for most of the remaining jurassic period. the jurassic succession on hold with hope is subdivided into shallow marine sandstones of the pelion and payer dal formations (vardekløft group) and lower shoreface – offshore transition heteroliths and offshore mudstones of the bernbjerg formation (hall bredning group). the pelion formation includes the new spath plateau member, which contains abundant sandy heteroliths in contrast to the dominant clean sandstone lithology of the pelion formation. the sedimentary facies of the units are described and the depositional environments interpreted. the jurassic succession is placed within a regional framework including the wollaston forland and jameson land basins to the north and south, respectively. ice undifferentiated superficial deposits dolerite sill plateau basalt normal fault inferred fault locality cross-section (log panel) paleocene lower cretaceous middle and upper jurassic lower triassic permian crystalline basement 1 8 9 3b 765 4 3a 3c 3d 3e 2 3 ■■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ steensby bjerg gael hamke bugt diener bjerg sortelv spath plateau fo sd al en stensiö plateau 21º00'w 21º15'w 74º00'n hold with hope g ul el v blåe lv ■ ■ ■■ ■■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ 2 km 73º55'n fig. 2. geological map of part of northern hold with hope including position of localities mentioned in text. the composite section from gulelv (fig. 5) was compiled from five part-sections (a–e). the solid lines show schematically the orientation of the ne–sw and nw–se log panels in figures 7 and 8, respectively. geus bulletin no 5.pmd 29-10-2004, 11:1453 54 geological setting the late palaeozoic – mesozoic extensional basin complex in east greenland is about 700 km long in a north– south direction. this complex is situated over structurally controlled en echelon troughs and forms a wedgeshaped embayment with the narrowest onshore part to the north. jurassic sediments are present in the wollaston forland and jameson land basins situated on the western margin of the rift complex. in both basins, sediment transport in middle jurassic time was mainly longitudinal from north to south along a low gradient basin floor, which was not differentiated into a shelf, slope and deep-water basin (surlyk 1977, 1990, 1991; surlyk et al. 1981; surlyk & clemmensen 1983). in the wollaston forland basin, rifting was initiated in middle jurassic time, and marine bajocian–bathonian sandstones onlap weathered caledonian basement rocks or upper permian carbonates. deposition took place on the hangingwall of wide w–sw-tilted fault blocks. jurassic rifting culminated in the volgian with strong rotational block faulting associated with conglomeratic submarine fan sedimentation (surlyk 1978). during this tectonic episode, the wide fault blocks originally defining the wollaston forland basin were split into smaller blocks (vischer 1943; surlyk 1978). the jurassic sediments on hold with hope occur on the hangingwall of small fault blocks that dip mainly towards the west and south-west (fig. 2). bedding planes within the triassic and jurassic successions seem tobeparallelwhereas theboundary to theoverlyingcretaceous succession is an angular unconformity (fig. 3). the jurassic succession shows marked lateral thickness variations depending on its position on the hangingwall. the thickness increases down-dip whereas it is missing up-dip due to early cretaceous erosion on some of the block crests. faults locally cut the triassic and jurassic successions, but not the overlying cretaceous succession showing that fault activity took place in post-kimmeridgian, but pre-barremian time (fig. 3). most faults in the area, however, were reactivated during cretaceous and cenozoic times. stratigraphy and sedimentology the jurassic succession on hold with hope is subdivided into lithostratigraphic units known from wollaston forland and kuhn ø (fig. 4). stratigraphic ages of the jurassic succession are based on ammonites and dinoflagellate cysts (piasecki et al. 2004, this volume). the oldest jurassic ammonite found on northern hold with hope is cranocephalites sp. indicating the upper bajocian c. pompeckji chronozone (j.h. callomon and p. alsen, personal communications 1997). the ammonite was, however, not found in situ but in the basal cretaceous conglomerate at locality 4 and c. 150 m east of locality 8 (fig. 2). the jurassic outcrops occur between stensiö plateau in the west and diener bjerg in the east (fig. 2). towards the south, the jurassic sediments are exposed along the eastern side of the gulelv river and on the southern side of the sortelv river. a composite section, 360 m thick, was measured in a north–south direction along gulelv on the hangingwall of a fault block dipping c. 15º towards the ssw and comprises segments 3a–e (figs 5, 6). lateral facies variations are illustrated by cross-sections oriented parallel (ne–sw) and perpendicular (nw–se) to the overall palaeocurrent direction (figs 2, 7, 8). pelion formation the pelion formation in the jameson land and wollaston forland basins consists of shallow marine, mediumto coarse-grained sandstones of late bajocian – late callovian age (engkilde & surlyk 2003). on hold pb e w lc mj tr tr fig. 3. triassic–paleocene succession exposed at steensby bjerg, northern hold with hope, viewed towards the south. the triassic (tr) and middle jurassic (mj) sediments dip towards the southwest and are unconformably overlain by lower cretaceous (lc) sediments. a post-kimmeridgian – pre-barremian normal fault striking north–south offsets the triassic and jurassic sediments by at least 30 m. view towards the south. from larsen et al. (1998). the exposure shown is c. 200 m high and capped by paleocene basalts (pb). geus bulletin no 5.pmd 29-10-2004, 11:1454 55 15 9. 4 15 4. 1 16 4. 4 16 9. 2 17 3 l l l l ba jo ci an ba th on ia n ju ra ss ic c al lo vi an o xf or di an k im m er id gi an st ag e sy st em m a m m m u u u u traill ø and geographical society ø w e w e s n hold with hope wollaston forland and kuhn ø be rn bj er g be rn bj er g o ly m pe n fo ss ilb j fb pp pe lio n br is to l e lv be rn bj er g pa ye r d al be rn bj er g ja ko bs st ig en pa ye r d al pe lio n ba st ia ns d al m us lin ge bj er g pe lio n sandstone siltstone heterolithic sandstone mudstone legend ammonite coal conglomerate c. cordatum p. baylei r. cymodoce a. mutabilis a. eudoxus a. autissiodorensis subboreal–boreal ammonite zonation c. densiplicatum c. tenuiserratum a. glosense a. serratum a. regulare a. rosenkrantzi p. plicatilis g. transversarium p. bifurcatus r. pseudocordata q. mariae q. lamberti p. athleta e. coronatum k. jason s. calloviense p. koenigi c. nordenskjoeldi m. herveyi c. discus o. orbis p. hodsoni m. morrisi t. subcontractus p. progracilis a. tenuiplicatus z. zigzag p. parkinsoni g. garantiana n. subfurcatum c. apertum c. calyx c. variabile a. cranocephaloide a. ishmae a. greenlandicus a. arcticus c. pompeckji c. indistinctus c. borealis fig. 4. jurassic lithoand biostratigraphy (lower jurassic and volgian not included) of the traill ø – geographical society ø region, hold with hope and wollaston forland. based on surlyk (1977, 1978, 1990, 1991), price & whitham (1997), alsgaard et al. (2003), alsen & surlyk (2004, this volume), vosgerau et al. (2004, this volume) and own observations in the traill ø and northern hold with hope areas. fossilbj/fb, fossilbjerget; pp, pelion (parnas mb). geus bulletin no 5.pmd 29-10-2004, 11:1455 56 lo w er s an ds to ne u ni t sp at h pl at ea u m b sp at h pl at ea u m b u gp ik r av in e m b m mud sand f c gr m mud sand f c gr m mud sand f c gr tr ia ss ic pe lio n fm pe lio n fm be rn bj er g fm be rn bj er g fm pa ye r d al f m c re ta ce ou s ? 120 110 100 90 80 70 60 50 40 30 20 10 0 250 240 230 220 210 200 190 180 170 160 150 140 130 360 350 340 330 260 m no exposure no exposure no exposure geus bulletin no 5.pmd 29-10-2004, 11:1456 57 with hope, a lower–middle callovian sandy succession, c. 190 m thick, overlying the lower triassic wordie creek formation is referred to the pelion formation (fig. 5). the succession is divided into two units. the lower unit is 30–40 m thick and consists of mediumto coarse-grained sandstones topped by silty, fine-grained sandstones. the upper unit is up to 155 m thick and differs from the clean sandstones that typify the formation by containing abundant sandy heteroliths interbedded with cross-bedded sandstones and is included in the spath plateau member (fig. 5). the lower sandstone unit and the spath plateau member are described and interpreted separately below; the latter member is defined formally as a new member of the pelion formation. structureless trough cross-bedding planar and trough cross-bedding structureless planar and trough cross-bedding wave ripple cross-lamination faint ripple form sets hummocky cross-stratification pebbles plant fragments log bivalve ammonite belemnite sedimentary structures biota miscellaneous features palaeocurrent direction orientation of wave ripple crest parallel lamination upper shoreface lower–middle shoreface facies associations offshore – lower shoreface offshore faint ripple form sets cross-lamination wave ripple cross-lamination curvolithos helminthopsis? chondrites bioturbation moderate bioturbation strong diplocraterion monocraterion ophiomorpha planolites bioturbation weak trace fossilsmudstone drapes and clasts fig. 5. composite sedimentary section of the middle–upper jurassic succession at gulelv (locality 3). the section was compiled from five part-sections located in a north–south direction along gulelv (figs 2, 6). modified from larsen et al. (1998). geus bulletin no 5.pmd 29-10-2004, 11:1457 58 lower sandstone unit the lower sandstone unit overlies siltstones and finegrained sandstones of the lower triassic wordie creek formation with a sharp erosional boundary. the upper boundary is placed where coarse-grained sandstones are overlain by silty, very fine-grained sandstones of the spath plateau member (figs 5, 9). the ammonite cadoceras cf. breve indicating the lower callovian c. apertum chronozone (j.h. callomon and p. alsen, personal communications 1997) was found 10 m above the triassic–jurassic boundary at stensiö plateau (fig. 2, locality 1). the basal part of the unit seems to be younger east of stensiö plateau (fig. 2, localities 4–9) where dinoflagellate cysts indicate the lower callovian p. koenigi chronozone. the ageof the upper boundary is constrained by an ammonite and by dinoflagellate cysts found in the basal part of the overlying spath plateau member indicating the lower callovian p. koenigi chronozone (see below). the lower sandstone unit consists of pebbly, mediumto coarse-grained quartz sandstones, which are trough cross-bedded with sets up to 0.5 m thick or locally appear structureless. small scour-fills with pebbles, up to 3 cm in diameter, are common. dip-directions of foresets show a dominance towards the sw, but directions towards the nw also occur (fig. 10). the sandstones commonly form coarsening-upwards units, 6–8 m thick, which locally are overlain by an erosionally based pebbly sandstone lag. the top part of the coarsening-upwards units is commonly calcite cemented and contains abundant vertical trace fossils of diplocraterion habichi. other trace fossils include monocraterion tentaculatum and locally ophiomorpha nodosa. belemnites, bivalves and silicified wood are common. the unit is slightly finer grained in places and medium-grained sandstones occur at locality 1, stensiö plateau (fig. 8). they contain low-relief scour surfaces draped by organic-rich mudstone and are strongly bioturbated with both horizontal and vertical burrows. the lower sandstone unit increases in thickness from c. 30 to 40 m from the sw towards the ne (fig. 7). deposition took place in a marine environment as reflected by marine macrofossils and trace fossils. the coarse-grained and pebbly sandstones and the dominance of vertical burrows suggest shallow-water deposition under high-energy conditions. the scour-fills are interpreted to have been formed by strong currents related to storm surges (e.g. clifton et al. 1971; hunter et al. 1979). the trough cross-bedded sandstones probably reflect 3-d dunes migrating seawards to the southwest in rip channels and partly longshore in runnels towards the north-west. erosionally-based pebbly sandstones with marine macrofossils form the uppermost beds of some of the coarsening-upwards units and are interpreted as marine lag deposits, formed by wave winnowing of underlying upper shoreface and foreshore deposits. the coarsening-upwards units are interpreted to have formed mainly by shallow marine shoreface progradation. the probable source of the calcite cement in the top part of the coarsening-up3d3c3b3a 3e payer dal fm bernbjerg fm ? 100 m lower triassic n s 100 m pelion and payer dal fms bernbjerg fm lower cretaceous unconformity pelion fm lower triassic spm lu fig. 6. sketch of triassic–cretaceous succession exposed in a fault block on the eastern side of the gulelv river (fig. 2). the location of the part-sections that make up the composite section in figure 5 are shown. the triassic and jurassic sediments dip c. 15° towards ssw and are unconformably overlain by cretaceous sediments, which dip c. 10° towards the south. an angular unconformity between the pelion and payer dal formations is observed on the north-facing valleyside at 3c (see fig. 11). lu, lower sandstone unit; spm, spath plateau member. geus bulletin no 5.pmd 29-10-2004, 11:1458 59 wards units is biogenic carbonate derived from calcareous shells accumulated on the marine omission surfaces (alsgaard et al. 2003; engkilde & surlyk 2003). the strong bioturbation and the thin mudstone layers in the medium-grained sandstones at stensiö plateau suggest deposition under lower energy conditions. the mudstone drapes on the scour-surfaces formed by suspension fall-out during fair-weather conditions following erosional storm events. the sandstones were probably deposited in a slightly deeper, middle shoreface environment, than the mediumto coarse-grained sandstones that dominate the lower sandstone unit east of stensiö plateau. the decrease in thickness of the sandstone unit from ne to sw combined with the dominant south-westwards palaeocurrent directions may reflect a ne–sw proximal–distal trend. in the proximal areas, a large sediment supply may have delayed the overall drowning recorded by the boundary to the overlying spath plateau member. the increase in thickness towards the ne may, however, also reflect the relief of the palaeoshoreface. spath plateau member new member general. the member comprises mainly cross-bedded sandstones and sandy heteroliths forming the upper part of the pelion formation on northern hold with hope. name. after the ice-covered basalt plateau south-west of gulelv, northern hold with hope. type locality. east side of gulelv (figs 2, 6, locality 3) where the composite type section (fig. 5) is defined from sub-sections a–c. thickness. 155 m. lithology. the spath plateau member consists of quartzitic sandstones and sandy heteroliths. a silty, very fine-grained sandstone bed occurs in the basal part of the member. belemnites, bivalves, silicified and coalified wood and impressions of leaf fragments are common. boundaries. the lower boundary is marked by an abrupt change from coarse-grained sandstone of the lower sandstone unit of the pelion formation to silty, fine-grained sandstone. the upper boundary is sharp and erosional and is overlain by pebbly sandstone and sandy heteroliths of the payer dal formation. distribution. the member occurs on northern hold with hope. age. early–late callovian based on an ammonite found at the base of the member and dinoflagellate cysts (fig. 7, locality 6). the ammonite is a microconch and resembles cadoceras septentrionale indicating the lowermost part of the lowercallovian p. koenigi chronozone (p. alsen, personal communication 1998). dinoflagellate cysts from the base of the member also indicate the c. apertum – p. koenigi chronozones (piasecki et al. 2004, this volume). dinoflagellate cysts in the upper part of the member indicate the upper callovian p.athletachronozone(piaseckietal.2004, thisvolume). facies description. the basal part of the member consists of dark brown, silty sandstones forming a 6 m thick marker bed, situated 30–40 m above the base of the jurassic succession (figs 5, 7, 9). the basal 0.5 m of the bed locally contain scattered fine pebbles up to 6 mm in diameter, but otherwise the unit coarsens upwards from silty, very fine-grained sandstone into silty fine-grained sandstone. the sandstones are horizontally laminated, structureless or show subtle wave or current cross-lamination with thin organic-rich mudstone drapes, but primary structures are to a large extent obscured due to weathering or strong bioturbation. in a few places, laminae of silty sandstone are deflected around small, elongate carbonate-cemented concretions. coalified wood fragments up to 30 cm long and belemnites are abundant in the basal part of the bed, and a few bivalves and gastropods were found immediately above the lower boundary at locality 6 (fig. 7) together with the ammonite cadoceras septentrionale. the organic content (toc) is low, about 1 wt%. the fine-grained marker bed forms the base of the spath plateau member and sharply overlies trough cross-bedded sandstones or conglomerates of the lower sandstone unit of the pelion formation. however, at locality 9 (fig. 8), the basal unit of the member consists of a ripple cross-laminated sandy heterolith unit, c. 0.4 m thick, followed by mediumto coarse-grained, planar cross-bedded sandstones. the upper boundary to the overlying ripple cross-laminated sandy heteroliths is gradational. geus bulletin no 5.pmd 29-10-2004, 11:1459 60 100 m 90 80 70 60 50 40 30 20 10 0 80 m 70 200 m 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 60 50 40 30 20 10 0 sand sand sand sand 0 10 20 30 40 50 60 70 80 m ? sw (distal) 2 3 4 5 ~ 2450 m ~ 1050 m ~ 130 m ~ 650 m lo w er u ni t sp at h pl at ea u m b pe lio n fm pa ye r d al f m geus bulletin no 5.pmd 29-10-2004, 11:1460 61 fig. 7. log panel giving a broadly ne (proximal) to sw (distal) cross-section. for legend, see fig. 5; for localities, see fig. 2. 10 0 20 30 40 50 60 70 m 40 m 30 20 10 0 10 0 60 m 50 40 30 20 sand sandsand t ri as si c ju ra ss ic s pa th p la te au m b c re ta ce ou s base of distributary channel fill deposits ne (proximal) 6 7 8 ~ 200 m ~ 1325 m geus bulletin no 5.pmd 29-10-2004, 11:1461 62 the bulk of the member is made up of heteroliths forming units up to 10 m thick of cross-laminated, fineto medium-grained sandstones with thin organic-rich mudstone drapes. both symmetrical and asymmetrical ripples, 1–5 cm high and with wavelengths up to 10 cm, occur. the heteroliths locally form cross-strata, up to 1.5 m thick, with very low-angle foresets and set boundaries marked by indistinct toesets of dark, sandy mudstones, a few centimetres thick. scour surfaces with a relief of up to 20 cm are common in the heteroliths. locally, structureless sandstones overlie the surfaces. the degree of bioturbation varies from moderate to high (up to 100%) and the trace fossil assemblage includes planolites beverleyensis, curvolithos multiplex, monocraterion tentaculatum, diplocraterion habichi and possibly helminthopsis magna. silicified and coalified wood fragments up to 1 m long, together with impressions of leaf fragments, are abundant and belemnites occur locally. two types of cross-bedded sandstones occur closely associated with the heteroliths in the spath plateau member. they both show foresets mainly dipping towards the south-west (fig. 10). the first type consists of planar or trough cross-bedded, fineto coarsegrained sandstones occasionally with pebbles. foresets are generally tangential and commonly separated by single and in places double, organic-rich mudstone drapes. backflow ripples and reactivation surfaces occur locally. the sets are 0.5–5 m thick and form cosets up to 25 m thick. set-boundaries are defined by organic-rich silty sandstone beds 1–10 cm thick, representing distal toesets. in one place, however, the toesets form a lenticular body of organic-rich shale, 40 m long and 0.5 m thick, with gently inclined laminae (to sw). the cross-bedded, coarse-grained sandstones commonly overlie ripple cross-laminated sandy heteroliths with a sharp erosional boundary, and contain belemnites, bivalve shells and coalified wood. the degree of bioturbation is low and the trace fossils include vertical burrows of monocraterion tentaculatum, diplocraterion habichi and arenicolites isp. the first type of cross-bedded sandstones form a c. 40 m thick succession in the lower part of the spath plateau member at locality 9, diener bjerg, whereas heteroliths are abundant in the lower part at the other localities (figs 7, 8). the second type of cross-bedded sandstones is fine0 10 20 30 40 50 60 70 80 m 20 m 10 0 10 0 20 30 40 50 60 70 m lo w er u ni t pe lio n fm t ri as si c ju ra ss ic c re ta ce ou s sp at h pl at ea u m b nw se sandsand sand 1 8 9~ 8 km ~ 5 km fig. 8. log panel giving a nw–se strike section through the jurassic succession. for legend, see fig. 5; for localities, see fig. 2. geus bulletin no 5.pmd 29-10-2004, 11:1462 63 to medium-grained and consists of up to 5 m thick sets of low-angle master beds separated by thinner crossbedded or ripple cross-laminated sets. the latter commonly show climbing ripple cross-lamination or locally bi-directional cross-laminae. the surfaces of the low-angle master beds may be wave or current rippled and separated by thin mudstone drapes. set thickness is 1–5 m. the lower boundary to ripple laminated sandy heteroliths is generally gradational. the sandstones contain belemnites, bivalve shells and coalified wood. the degree of bioturbation is moderate to high. trace fossils include vertical burrows of monocraterion tentaculatum, diplocraterion habichi and arenicolites isp., forms that are also common in the first type of sandstones, together with horizontal burrows of planolites beverleyensis and possibly helminthopsis magna in the intervening mud drapes. trough cross-bedded sandstones similar to those in the lower sandstone unit of the pelion formation occur locally in the spath plateau member (fig. 7). they commonly form coarsening-upwards successions, up to 5 m thick, the most complete of which have a lower part consisting of well-sorted, fineto medium-grained sandstone with indistinct ripple cross-lamination, trough cross-bedding and shallow scour fills. the scour fills are 10–30 cm thick and are marked by thin, organicrich mudstone layers above the lower erosional boundaries followed by laminated or structureless sandstones. the upper part of the coarsening-upwards units consists of trough cross-bedded, mediumto coarse-grained sandstones. foresets dip mainly towards the south, but dip-directions towards the west and east also occur (fig. 10). the upper part is commonly calcite cemented and capped by a sharp surface from which abundant diplocraterion habichi descend; other trace fossils in the upper levels include monocraterion tentaculatum and occasional ophiomorpha nodosa. horizontal burrows of planolites beverleyensis are limited to the lower part of the coarsening-upwards units. belemnites, bivalves and silicified wood are common. facies interpretation. the laminated silty sandstone 5 m spm lu pelion fm tr ias sic fig. 9. triassic siltstones and sandstones overlain by concretionary sandstones of the lower sandstone unit (lu) of the pelion formation. a marine drowning surface (dashed line) separates the lower sandstone unit from an overlying dark brown silty sandstone bed which form the basal part of the spath plateau member (spm), pelion formation. fig. 2, locality 6. geus bulletin no 5.pmd 29-10-2004, 11:1463 64 bed of the basal spath plateau member (except at locality 9), was deposited in a marine environment as reflected by the marine dinoflagellate cysts and abundant belemnites. the horizontal lamination was formed by deposition from suspension fall-out whereas the subtle cross-lamination was probably formed by waveinduced currents in the offshore transition to lower shoreface zone. the sharp boundary to the underlying upper shoreface sandstones of the lower sandstone unit represents a drowning surface. at locality 9, diener bjerg, this surface is overlain by a sandy heterolith unit, c. 0.4 m thick, interpreted as having been deposited in the lower to middle shoreface zone. this lateral facies variation suggests that the palaeo-water depth decreased from west to east. the alternation of ripple cross-laminated sandstones 25 20 15 10 5 % 50 40 30 20 10 5 % 50 40 30 20 10 5 % 25 1520 10 5 % 25 20 15 10 5 % 15 10 5 % n=11 v=243° n=9 v=221° n=11 v=175° n=14 v=338° n=47 v=220° n=8 v=231° pelion fm (lower sandstone unit) pelion fm (spath plateau member) trough cross-bedding trough and planar cross-bedding trough cross-bedding wave ripple crest orientation trough and planar crossbedding, and low-angle master bedding trough cross-bedding payer dal fm bernbjerg fm fig. 10. palaeocurrent and wave-ripple data from the pelion, payer dal and bernbjerg formations. rose diagrams are shown as true area plots. v, vector mean; n, number of measurements. geus bulletin no 5.pmd 29-10-2004, 11:1464 65 and mudstone drapes reflects varying energy conditions and may be related to the alternation of fairweather and storm-wave processes and perhaps tidal currents. the strongly bioturbated heteroliths reflect periods of slow sedimentation and little physical reworking whereas the scour surfaces and the overlying structureless sandstones which form part of the facies are interpreted as having formed by strong currents related to storm surges (clifton et al. 1971; hunter et al. 1979; nemec & steel 1984). the abundant plant debris suggests a significant sediment supply from land. the two types of cross-bedded sandstones are interpreted to represent ssw-migrating sandwaves. they were modified by waves and opposing tidal currents as seen by the presence of wave-ripples, mud drapes, reactivation surfaces and bimodal cross-lamination (visser 1980; boersma & terwindt 1981; wood & hopkins 1989). the first type of cross-bedded sandstones was formed in shallow water, possibly the upper shoreface, as seen by the coarse grain-size and low degree of bioturbation. the sharp erosional lower boundaries to lower–middle shoreface heteroliths suggest that the migration of sandwaves took place in distributary channels. the marked lateral facies variation seen between locality 9 and the other localities suggests that the thick cross-bedded succession at locality 9 may represent a major distributary channel fill. the second type of cross-bedded sandstones was formed in deeper water than the first type, as reflected by the generally finer grain-size, stronger bioturbation and the gradational boundary to underlying lower– middle shorefaceheteroliths.the compound cross-bedded sandstones with climbing ripple cross-laminated sets on the low-angle master bedding reflect suspension fall-out into deeper water and were possibly deposited in mouthbars (elliott 1974; gjelberg & steel 1995). the trough cross-bedded sandstones are similar to those in the lower sandstone unit of the pelion formation and are similarly interpreted as having formed by migration of 3-d dunes in a high energy, shallow marine environment. the coarse grain size of the trough cross-beds from the upper part of the coarsening-upwards units suggests sediment supply from nearby distributary channels. the dominant southwards palaeocurrent direction indicates that land was situated towards the north, but the large variation in the palaeocurrent measurements and the limited dataset preclude detailed interpretation of the palaeo-shoreline orientation. the coarsening-upwards units are interpreted to reflect progradation in the middle to upper shoreface. payer dal formation in the type area on kuhn ø, the payer dal formation consists of fineto coarse-grained quartz sandstones locally with pebbly sandstone lags rich in marine bivalves and belemnites (alsgaard et al. 2003). the formation is subdivided into a lower and an upper unit that have an early–middle oxfordian and early late oxfordian age, respectively (fig. 4). on hold with hope, a succession of cross-bedded, mediumto coarse-grained quartz sandstones, pebbly sandstone lags and sandy heteroliths overlying the pelion formation is referred to the payer dal formation. e w lower cretaceous payer dal fm pelion fm(spath plateau mb) 15 m fig. 11. the spath plateau member (pelion formation), payer dal formation and overlying cretaceous sandstones. note the slight angular discordance at the unconformity between the pelion and payer dal formations (dashed line). fig. 2, locality 3c; view towards south. geus bulletin no 5.pmd 29-10-2004, 11:1465 66 the age of the payer dal formation on hold with hope is middle–late oxfordian. dinoflagellate cysts from the lower part of the formation indicate the middle–upper oxfordian c. tenuiserratum – a. glosense chronozones, similar to the upper unit of the payer dal formation on kuhn ø. this suggests that the boundary between the pelion and payer dal formations on hold with hope represents a late callovian – middle oxfordian hiatus (fig. 4). the formation is exposed at localities 2 and 3 (figs 2, 5, 7). the lower boundary is only exposed at locality 3, where it is represented by an erosional surface forming a minor angular unconformity between the pelion and payer dal formations (fig. 11). the surface is overlain by a pebbly sandstone lag, up to 0.5 m thick, of quartz pebbles 3–5 mm in diameter, belemnites, bivalves and small logs, up to 0.4 m long. the lag is overlain by a coarseto very coarse-grained sandstone succession, c. 15 m thick, which fines slightly upwards. the succession is dominated by trough crossbedding but small sets of planar cross-beds and pebbly sandstone lags also occur. dip directions of foresets are mainly towards the sw (fig. 10). pavements of bivalves are common, whereas belemnites, logs and rounded mudstone clasts, up to 5 cm in diameter, occur locally. the sandstone succession is capped by a sharp surface overlain by an overall coarsening-upwards unit (c. 25 m thick) of ripple cross-laminated sandy heteroliths, similar to those in the underlying spath plateau member. at locality 2, the formation is at least 70 m thick and consists of alternating ripple cross-laminated sandy heteroliths and sets of planar or trough cross-bedded, coarse-grained sandstones similar to the first type of cross-bedded sandstones of the spath plateau member. the foresets dip towards the south-west (fig. 10). the upper boundary to the bernbjerg formation is covered by scree, but is probably situated somewhere between 223 m and 234 m in the composite section at locality 3 (fig. 5). the basal pebbly sandstone at locality 3 is interpreted as a marine lag deposit due to the coarse grain size, the presence of marine macrofossils and the erosional base. the overlying cross-bedded sandstone succession was deposited in a high energy, shallow marine environment as testified by the coarse grain size, the pebbly sandstone lags and the marine macrofossils. the cross-bedded sandstones are interpreted to reflect the seawards migration of dunes in the upper shoreface. the coarse grain size of the sandstone succession and the occurrence of logs and rounded clay clasts suggest sediment supply from nearby distributary channels. the sandstone succession is topped by a drowning surface. a slightly different depositional environment is recorded by the succession at locality 2 where planar and trough cross-bedded distributary channel-fill sandstones alternate with lower–middle shoreface heteroliths. these facies are very similar to the facies of the underlying spath plateau member and are similarly interpreted to reflect migration of dunes in distributary channels and mouth bars associated with tidally influenced deltas (see above). the considerable thickness variation of the payer dal formation between localities 2 and 3, together cretaceous bernbjerg fm fig. 12. structureless and horizontally laminated dark mudstones of the bernbjerg formation erosionally overlain by cretaceous sandstone (fig. 6, locality 3e; 340–360 m in fig. 5). person (encircled) for scale. geus bulletin no 5.pmd 29-10-2004, 11:1466 67 with the presence of a minor angular unconformity at the base of the formation, might reflect differential subsidence due to the onset of fault-block tilting. bernbjerg formation the upper oxfordian – lower volgian bernbjerg formation in wollaston forland is dominated by darkgrey to black mudstones and strongly bioturbated heteroliths (surlyk 1977; surlyk & clemmensen 1983). on hold with hope, the formation is poorly exposed along the gulelv river (figs 2, 6) where it is estimated to be c. 130 m thick based mainly on simple geometrical calculations. it has not been possible to study lateral facies variations within the bernbjerg formation. the basal part of the bernbjerg formation on hold with hope has a late oxfordian, a. glosense chron age based on the presence of the ammonite amoeboceras ilovaiskii (j.h. callomon and p. alsen, personal communications 1997) and dinoflagellate cysts (piasecki et al. 2004, this volume). in the upper part of the formation, dinoflagellate cysts indicate a late oxfordian –earlykimmeridgian, a. serratum – p. baylei chronage (piasecki et al. 2004, this volume). the basal part of the bernbjerg formation on hold with hope is dominated by organic-rich, silty, finegrained sandstones, which are horizontally laminated or locally cross-laminated. wave-rippled, mediumto coarse-grained sandstone beds, locally with pebbles up to 1 cm in diameter, are common. the wave ripples have nw–se ripple crest orientations (fig. 10; 234– 248 m in fig. 5). the silty fine-grained sandstones contain scattered chondrites isp. belemnites occur locally whereas ammonites are abundant. this lower coarsergrained unit of the bernbjerg formation is referred to the ugpik ravine member (surlyk 2003, fig. 5). the succession above the basal part of the bernbjerg formation consists of structureless or horizontally laminated dark mudstones (fig. 12). no macrofossils were found in the mudstones. total organic carbon (toc) content is about 2 wt% based on two samples. the lower boundary of the bernbjerg formation is not exposed, and the formation is erosionally overlain by lower cretaceous coarse-grained sandstones. the bernbjerg formation probably continues into the subsurface along gulelv. the horizontally laminated and locally cross-laminated silty fine-grained sandstones from the basal part of the bernbjerg formation reflect deposition from suspension fall-out and weak bottom currents during fair-weather conditions. storm-wave currents most likely deposited the interbedded wave-rippled, medium to coarse-grained sandstones. the orientations of the wave ripples suggest a nw–se-trending coastline. the facies are interpreted to have been deposited in the offshore transition to lower shoreface zone. the overlying dark mudstones are interpreted to have been deposited offshore from suspension fall-out based on the fine grain size and dominant horizontal lamination. palaeoenvironments and basin configuration the jurassic succession on hold with hope shows a stepwise, but overall fining-upwards trend (fig. 5) reflecting long-term transgression. the transgression was interrupted by a major relative sea-level fall within the late callovian – middle oxfordian time interval, represented by the unconformity separating the pelion and payer dal formations. the basal jurassic sediments seem to get younger from west to east as suggested by dinoflagellate cysts indicating the lower callovian c. apertum chronozone at stensiö plateau and the lower callovian p. koenigi chronozone at steensby bjerg and diener bjerg (fig. 2). the age difference may reflect onset of rifting and associated onlap. alternatively, the stensiö plateau formed a separate fault block, which was transgressed first. the n–s-trending fault situated at blåelv, west of stensiö plateau, forms the western boundary of jurassic sediments today (fig. 2). the jurassic sea, however, may have extended as far west as to the postdevonian main fault, which formed the western basin boundary during the late permian and triassic (vischer 1943; birkelund & perch-nielsen 1976; stemmerik et al. 1993). this would imply that older jurassic sediments from the western part of the basin were eroded in post-kimmeridgian time. transport of eroded material towards the east may explain the occurrence of cranocephalites sp., indicating the upper bajocian c. pompeckji chronozone, in the cretaceous basal conglomerate at steensby bjerg. the sediment source area may, however, also have been clavering ø, west of the dombjerg–clavering fault, which is believed to have formed a palaeo-high in jurassic times (vischer 1943; stemmerik et al. 1993). lateral facies variations within the spath plateau member (pelion formation) indicate that palaeo-water depths decreased from west to east. this suggests a similar setting to that envisaged for the wollaston forgeus bulletin no 5.pmd 29-10-2004, 11:1467 68 land basin where deposition took place on the w– sw-tilted hangingwalls of major fault blocks and the elevated fault block crests formed elongated islands or peninsulas to the east (vischer 1943; maync 1947; donovan 1957; surlyk 1977; surlyk et al. 1981; surlyk & clemmensen 1983). the presence of the marine pelion and payer dal formations in the vicinity of the block crest excludes the occurrence of a major land area during callovian and middle oxfordian time when the crest probably only formed elongated islands or submarine shoals. the hold with hope area is thus interpreted to have formed a narrow embayment which was open for marine circulation towards the south and, during periods of high sea level, eastwards across the elevated fault block crest. farther towards the east, the fault block was most likely limited by the continuation of the dombjerg–clavering fault, which was active during the jurassic (maync 1947; surlyk 1977; stemmerik et al. 1993). the narrow head of the rift-basin occurs in an intermediate position between the wollaston forland and jameson land basins that are situated to 100 km 20ºw 16ºw24ºw28ºw 76ºn 74ºn 72ºn wollaston forland basin hold with hope traill ø jameson land basin upper shoreface lower shoreface offshore inferred coastline main direction of sediment transport early–middle callovian fig. 13. early–middle callovian palaeogeography and facies distribution in east greenland. based on surlyk (1977), engkilde & surlyk (2003), vosgerau et al. (2004, this volume), and new data from hold with hope. geus bulletin no 5.pmd 29-10-2004, 11:1468 69 the north-east and south-west, respectively (fig. 13). the hold with hope region seems to have formed a land area during the late callovian – middle oxfordian time interval whereas it was covered by sea during maximum flooding in late bajocian, early–middle callovian and late oxfordian – kimmeridgian times. a close comparison of the jurassic successions in the wollaston forland, hold with hope and jameson land basins is hindered by limited biostratigraphic control at some levels due to the scarcity of ammonites and dinoflagellate cysts in the coarse sandstone facies. it is evident, however, that the late callovian – middle oxfordian hiatus in the hold with hope basin has not been demonstrated in the other two basins where sediments of early–middle oxfordian age are well documented (fig. 4). the minor angular unconformity between the pelion and payer dal formations suggests that the hiatus was formed by minor tilting of fault blocks and erosion of uplifted block crests perhaps combined with eustatic sea-level falls. sea-level falls, possibly eustatic, have been suggested to take place in late callovian time, at the early–middle oxfordian boundary and in early late oxfordian time (sahagian et al. 1996) and seem to correspond to changes in regional sea level documented in jameson land (larsen & surlyk 2003). the dark mudstones of the bernbjerg formation reflect deposition in a quiet offshore environment indicating that the influence of basin topography was overprinted by rise in relative sea level. fault block crests to the east were finally inundated during the kimmeridgian. strong fault activity occurred at the boundary between the post-kimmeridgian and pre-barremian successions and the fault blocks originally defining the hold with hope basin were split into smaller blocks, a similar tectonic development to that seen in the wollaston forland basin. comparison with the wollaston forland basin suggests that this tectonic episode took place in volgian–valanginian time. summary and conclusions a new middle–upper jurassic succession is described from northern hold with hope. the jurassic sediments occur on the hangingwall of small fault blocks dipping towards the wsw. the sediments are locally eroded away on the uplifted block crests, whereas they increase in thickness down-dip on the hangingwalls. the jurassic succession is up to 360 m thick and is referred to the pelion, payer dal and bernbjerg formations of early callovian – early kimmeridgian age. it overlies lower triassic siltstones and sandstones with a sharp boundary and is overlain by cretaceous coarse-grained sandstones with an angular unconformity. the succession was deposited in a shoreface–offshore marine environment and reflects an overall transgression. the pelion formation is c. 190 m thick and consists of a lower sandstone unit, 30–40 m thick, overlain by a drowning surface and sandy heteroliths and sandstones of the new spath plateau member, c. 155 m thick. the lower sandstone unit spans the lower callovian c. apertum – p. koenigi chronozones. it consists of medium to coarse-grained sandstones, which are trough cross-bedded or structureless and contain small pebbly scour-fills. the sandstones commonly form coarsening-upwards units, 6–8 m thick, the upper part of which are calcite cemented and contain abundant diplocraterion habichi. the sandstones are interpreted to have been deposited in the middle to upper shoreface zone. the spath plateau member is of late early – middle callovian, p. koenigi – p. athleta chron age based on ammonites and dinoflagellate cysts. it is dominated by lower–middle shoreface ripple cross-laminated sandy heteroliths and cross-bedded sandstones reflecting migration of 2-d and 3-d dunes in distributary channels and mouth bars associated with tidally influenced deltas. the payer dal formation is more than 70 m thick and of middle–late oxfordian, c. tenuiserratum – a. glosense chron age. it consists of cross-bedded sandstones and sandy heteroliths showing considerable lateral thickness variations and reflects progradation of tidally influenced dunes and mouth bars. it is separated from the underlying pelion formation by an angular unconformity representing a late callovian – middle oxfordian hiatus, formed by subaerial erosion related to an eustatic sea-level fall, combined with minor tiltingof fault blocks and erosion of uplifted block crests. the bernbjerg formation is at least 130 m thick and spans the upper oxfordian – lower kimmeridgian (a. glosense – p. baylei chronozones). the basal part consists of horizontally laminated and locally cross-laminated silty fine-grained sandstones interbedded with thin wave-rippled mediumto coarse-grained sandstones and is referred to the ugpik ravine member. deposition took place in the offshore transition to lower shoreface zone. the upper part of the formation consists of structureless or horizontally laminated dark mudstones reflecting offshore deposition in an outer shelf environment. geus bulletin no 5.pmd 29-10-2004, 11:1469 70 deposition of the jurassic succession on hold with hope took place in a narrow rift-controlled embayment as indicated by the distribution of the sediments and dominating south-westwards palaeocurrent directions in the pelion and payer dal formations. lateral facies variations in the spath plateau member indicate a decrease in palaeowater depth from west to east. the embayment was open to marine circulation to the south and probably partly to the east, where the crestal margin of a slightly westwards to south-westwards tilted block formed elongated narrow islands or submarine shoals. during the deposition of the bernbjerg formation, the influence of basin topography was overprinted by a rise in relative sea level and the fault block crest to the east was inundated. towards the west, the stensiö plateau may have formed a separate block that was transgressed first, as indicated by the fact that the pelion formation is older at this locality than at steensby bjerg and diener bjerg. the presence of an angular unconformity between the jurassic and cretaceous succession shows that block rotation took place in post-kimmeridgian – pre-barremian time. during this tectonic episode, the fault block originally defining the hold with hope basin was split into narrower blocks. this might have taken place in the volgian–valanginian as suggested by comparison with the wollaston forland basin where a similar tectonic event took place during this time interval. acknowledgements the present study received support from saga petroleum asa, and is a contribution to the project ‘resources of the sedimentary basins of north and east greenland’ supported by the danish research councils. we are grateful to lars stemmerik and the referees jon gjelberg and finn surlyk for very useful and constructive comments. peter alsen and john h. callomon are thanked for providing biostratigraphic information. references alsen, p. & surlyk, f. 2004: maximum middle jurassic transgression in east greenland: evidence from new ammonite finds, bjørnedal, traill ø. in: stemmerik, l. & stouge, s. (eds): the jurassic of north-east greenland. geological survey of denmark and greenland bulletin 5, 31–49 (this volume). alsgaard, p.c., felt, v.l., vosgerau, h. & surlyk, f. 2003: the jurassic of kuhn ø, north-east greenland. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 865–892. birkelund, t. & perch-nielsen, k. 1976: late palaeozoic – mesozoic evolution of central east greenland. in: escher, a. & watt, w.s. (eds): geology of greenland, 304–339. copenhagen: geological survey of greenland. boersma, j.r. & terwindt, j.h.j. 1981: neap-spring tide sequences of intertidal shoal deposits in a mesotidal estuary. sedimentology 28, 151–170. clifton, h.e., hunter, r.e. & phillips, r.l. 1971: depositional structures and processes in the non-barred high-energy nearshore. journal of sedimentary petrology 41, 651–670. donovan, d.t. 1957: the jurassic and cretaceous systems in east greenland. meddelelser om grønland 155(4), 214 pp. elliott, t. 1974: interdistributary bay sequences and their genesis. sedimentology 21, 611–622. engkilde, m. & surlyk, f. 2003: shallow marine syn-rift sedimentation: middle jurassic pelion formation, jameson land, east greenland. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 813–863. gjelberg, j. & steel, r.j. 1995: helvetiafjellet formation (barremian–aptian), spitsbergen: characteristics of a transgressive succession. in: steel, r.j., felt, v., johannessen, e.p. & mathieu, c. (eds): sequence stratigraphy on the northwest european margin. norwegian petroleum society (npf) special publication 5, 571–593. amsterdam: elsevier. hunter, r.e., clifton, h.e. & phillips, r.l. 1979: depositional processes, sedimentary structures and predicted vertical sequences in barred nearshore systems, southern oregon coast. journal of sedimentary petrology 49, 711–726. kelly, s.r.a., whitham, a.g., koraini, a.m. & price, s.p. 1998: lithostratigraphy of the cretaceous (barremian–santonian) hold with hope group, ne greenland. journal of the geological society of london 155, 993–1008. koch, l. 1932: carboniferous and triassic stratigraphy of east greenland. meddelelser om grønland 83(2), 100 pp. koch, l. & haller, j. 1971: geological map of east greenland 72°–76°n. lat. (1:250 000). meddelelser om grønland 183, 26 pp., 13 maps. larsen, m. & surlyk, f. 2003: shelf-edge delta and slope deposition in the upper callovian – middle oxfordian olympen formation, east greenland. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 931–948. larsen, m., piasecki, s., preuss, t., seidler, l., stemmerik, l., therkelsen, j. & vosgerau, h. 1998: petroleum geological activities in east greenland in 1997. geology of greenland survey bulletin 180, 35–42. maync, w. 1947: stratigraphie der jurabildungen ostgrönlands zwischen hochstetterbugten (75°n) und dem kejser frantz joseph fjord (73°n). meddelelser om grønland 132(2), 223 pp. maync, w. 1949: the cretaceous beds between kuhn island and cape franklin (gauss peninsula), northern east greenland. meddelelser om grønland 133(3), 291 pp. nemec, w. & steel, r.j. 1984: alluvial and coastal conglomerates: their significant features and some comments on gravgeus bulletin no 5.pmd 29-10-2004, 11:1470 71 elly mass-flow deposits. in: koster, e.h. & steel, r.j. (eds): sedimentology of gravels and conglomerates. canadian society of petroleum geologists memoir 10, 1–31. piasecki, s., larsen, m., therkelsen, j. & vosgerau, h. 2004: jurassic dinoflagellate cyst stratigraphy of hold with hope, north-east greenland. in: stemmerik, l. & stouge, s. (eds): the jurassic of north-east greenland. geological survey of denmark and greenland bulletin 5, 73–88 (this volume). price, s.p. & whitham, a.g. 1997: exhumed hydrocarbon traps in east greenland: analogs for the lower–middle jurassic play of northwest europe. american association of petroleum geologists bulletin 81, 196–221. sahagian, d., pinous, o., olferiev, a. & zakharov, v. 1996: eustatic curve for the middle jurassic – cretaceous based on russian platform and siberian stratigraphy: zonal resolution. american association of petroleum geologists bulletin 80, 1433–1458. stemmerik, l., christiansen, f.g., piasecki, s., jordt, b., marcussen, c. & nøhr-hansen, h. 1993: depositional history and petroleum geology of the carboniferous to cretaceous sediments in the northern part of east greenland. in: vorren, t.o., bergsager, e., dahl-stamnes, ø.a., holter, e., johansen, b., lie, e. & lund, t.b. (eds): arctic geology and petroleum potential. norwegian petroleum society (npf) special publication 5, 67–68. amsterdam: elsevier. stemmerik, l., clausen, o.r., korstgård, j., larsen, m., piasecki, s., seidler, l., surlyk, f. & therkelsen, j. 1997: petroleum geological investigations in east greenland: project ‘resources of the sedimentary basins of north and east greenland’. geology of greenland survey bulletin 176, 29–38. surlyk, f. 1977: stratigraphy, tectonics and palaeography of the jurassic sediments of the areas north of kong oscar fjord, east greenland. bulletin grønlands geologiske undersøgelse 123, 56 pp. surlyk, f. 1978: submarine fan sedimentation along fault scarps on tilted fault blocks (jurassic–cretaceous boundary, east greenland). bulletin grønlands geologiske undersøgelse 128, 108 pp. surlyk, f. 1990: timing, style and sedimentary evolution of late palaeozoic – mesozoic extensional basins of east greenland. in: hardman r.f.p. & brooks, j. (eds): tectonic events responsible for britain’s oil and gas reserves. geological society special publication (london) 55, 107–125. surlyk, f. 1991: sequence stratigraphy of the jurassic – lowermost cretaceous of east greenland. american association of petroleum geologists bulletin 75, 1468–1488. surlyk, f. 2003: the jurassic of east greenland: a sedimentary record of thermal subsidence, onset and culmination of rifting. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 659–722. surlyk, f. & clemmensen, l.b. 1983: rift propagation and eustacy as controlling factors during jurassic inshore and shelf sedimentation in northern east greenland. sedimentary geology 34, 119–143. surlyk, f., callomon, j.h., bromley, r.g. & birkelund, t. 1973: stratigraphy of the jurassic – lower cretaceous sediments of jameson land and scoresby land, east greenland. bulletin grønlands geologiske undersøgelse 105, 76 pp. surlyk, f., clemmensen, l.b. & larsen, h.c. 1981: post-paleozoic evolution of the east greenland continental margin. in: kerr, j.w. & fergusson, a.j. (eds): geology of the north atlantic borderlands. canadian society of petroleum geologists memoir 7, 611–645. vischer, a. 1943: die postdevonische tektonik von ostgrönland zwischen 74° und 75°n br. kuhn ø, wollaston forland, clavering ø und angrenzende gebiete. meddelelser om grønland 133(1), 195 pp. visser, m.j. 1980: neap-spring cycles reflected in holocene subtidal large-scale bedform deposits: a preliminary note. geology 8, 543–546. vosgerau, h., alsen, p., carr, i.d., therkelsen, j., stemmerik, l. & surlyk, f. 2004: jurassic syn-rift sedimentation on a seawards-tilted fault block, traill ø, north-east greenland. in: stemmerik, l. & stouge, s. (eds): the jurassic of north-east greenland. geological survey of denmark and greenland bulletin 5, 9–18 (this volume). wood, j.m. & hopkins, j.c. 1989: reservoir sandstone bodies in estuarine valley fill: lower cretaceous glauconitic member, little bow field, alberta, canada. american association of petroleum geologists bulletin 73, 1361–1382. geus bulletin no 5.pmd 29-10-2004, 11:1471 geological survey of denmark and greenland bulletin 35, 2016, 67-70 67© 2016 geus. geological survey of denmark and greenland bulletin 35, 67–70. open access: www.geus.dk/publications/bull the marine record of the independence–danmark fjord system extending out to the wandel hav in eastern north greenland (fig. 1a) is little known due to the almost perennial sea-ice cover, which makes the region inaccessible for research vessels (nørgaard-pedersen et al. 2008), and only a few depth measurements have been conducted in the area. in 2015, the villum research station, a new logistic base for scientific investigations, was opened at station nord. in contrast to the early exploration of the region, it is now possible to observe and track the seasonal character and changes of ice in the fjord system and the arctic ocean through remote sensing by satellite radar systems. satellite data going back to the early 1980s show that the outer part of the independence–danmark fjord system is characterised by perennial sea ice whereas both the southern part of the fjord system and an area 20–30 km west of station nord are partly ice free during late summer (fig. 1b). hence, marine-orientated field work can be conducted from the sea ice using snow mobiles, and by drilling through the ice to reach the underlying water and sea bottom. earlier studies have shown that the last deglaciation of the region occurred in the early holocene (funder 1989; nørgaard-pedersen et al. 2008) and the sea subsequently inundated the fjord system. based on onshore evidence from eastern north greenland, beach ridges and frequent deposition of drift wood during the holocene thermal maximum at about 8000–5000 years bp indicate a period of open water and a mean summer temperature higher than today (funder et al. 2011). during the first field season in 2015 at the new station, a number of field-based research projects were carried out coordinated by the arctic science partnership (asp, http:// www.asp-net.org/) and the arctic research centre (arc) at aarhus university, denmark. the geological survey of denmark and greenland (geus) and arc carried out the fjord sediment coring project described here. the main objective of the marine geoscience field work was to collect sediment cores, which can be used to reconstruct past sea-ice variability through recent centuries and millennia. a major aspect was also the impact of sea ice on primary production, mainly diatoms and dinoflagellates and on biogeochemical cycles and the arctic ecosystem. this paper presents preliminary results of the sediment coring work and also new information on the bathymetry in the fjord area up to 50 km from station nord. a time-series of satellite radar images of sea-ice types in the fjord system provided by the danish meteorological institute (dmi) for the last decade is discussed to better understand the recent sea-ice dynamics in the area. an introduction to the ongoing laboratory work and data processing is also included. niels nørgaard-pedersen, sofia ribeiro, naja mikkelsen, audrey limoges and marit-solveig seidenkrantz investigations of past climate and sea-ice variability in the fjord area by station nord, eastern north greenland fig. 1. a: study area (framed area) near station nord and villum research station (red star), eastern north greenland. the average position of the fast-ice edge toward the arctic ocean pack-ice drift is indicated. map source: ibcao vers. 3.0 (jakobsson et al. 2012). b: satellite image (dmi aqua) of the area from 15 august 2015. pdø: prinsesse dagmar ø. ptø: prinsesse thyra ø. pmø: prinsesse margrethe ø. arctic ocean flade isblink 50 km station nord n pmø ptø pdø independence fjord hagen fjord danm ark fjo rd arctic ocean wandel hav flade isblink 50 km station nord pdø pmøptø n fast ice margin a b g re en la n d 6868 ice conditions satellite data show that the northern hemisphere has experienced a significant decrease in sea ice during the past c. 35 years, with a late summer sea-ice net loss in excess of 10% per decade, in addition to decreases seen in ice extent, thickness and age (perovich & richter-menge 2015). older (>4 years) and thicker sea ice is now only dominant in northern arctic canada and off northern greenland. the large independence–danmark fjord system in eastern north greenland is characterised by a semi-permanent fast-ice cover (wadhams 1986) in the outer part and in the inner more southern part by a seasonal ice cover, which breaks up and partly melts during august–september (fig. 1a). a coastal open water lead (polynya) is often observed between the northern margin of the fast ice and the drifting arctic pack ice. the fast ice in the outer fjord system is protected by three large islands (the ‘prinsesse islands’; figs 1, 2), which make it stable for a number of years before parts of it break up (wadhams 1986 and available satellite data). very thick and old ice of the ‘sikkussak’ type has, therefore, been observed earlier at the mouth of the fjord system (wadhams 1986). during exceptionally warm summers the outer fast-ice cover has broken up, for example as observed in mid-august 2003 (rasmussen 2004). the 2003 event coincided with break-up of the thick fast ice off the large ne greenland ice stream in 2002 and 2003. while planning our field work we used synthetic aperture radar (sar) and visible band images (aqua, terra) captured over the last decade from dmi (http:// ocean.dmi.dk/arctic/nord.uk.php) in order to characterise the multi-year ice dynamics of the fjord system and to select the sampling sites. sea-ice radar reflectivity is sensitive to the roughness of the ice and the presence of saltwater droplets within newer ice. thus, older and more deformed multi-year ice appears white or light grey (more reflection), whereas younger, first-year ice appears dark grey or black (less reflection). hence, it is possible to identify areas characterised by first-year ice and distinguish these from multiyear ice covered areas (fig. 2). moreover, by examining the ice-cover character on sar images from dmi from 2009 onwards, it is possible to date different patches of fast-ice cover. ice tongues of glacier ice and icebergs debouching out to the sea from the ice cap of flade isblink can be identified north-east and south-west of station nord (fig. 2). the sar and visible band image record for the past decade reveals that the oldest fast ice in the study area is found k22, 23 k33, 34 k1, 2, 3 k11, 12 k29, 30 k4, 5 k27, 28 k7, 8 k26 k31, 32 r1, 2 k9, 10 k17, 18, 19 k24, 25 k15, 16 k13, 14 k20, 21 depth (m) 0–20 20–40 40–60 60–80 80–100 100–120 120–140 140–160 160–180 180–200 10 km 81°50´n 81°30´n 81°30´n 16°w17°w prinsesse dagmar ø flade isblink station nord a b prinsesse thyra ø prinsesse dagmar ø multi-year ice first-year ice first-year ice glacier front glacier front station nord flade isblink 10 km 81°50´n 81°50´n 16°w17°w k33, 34 k1, 2, 3 k11, 12 k29, 30 k4, 5 k27, 28 k7, 8 k26 k31, 32 r1, 2 k9, 10 k17, 18, 19 k24, 25 k15, 16 k13, 14 k20, 21 k22, 23 18°w 18°w 81°40´n 81°40´n fig. 2. a: radar satellite image of the study area (sar sentinel-1 from 29.01.15), showing areas of seasonal (dark grey) and multi-year ice (light grey). there is no open water. high resolution sar image (courtesy of leif toudal, danish meteorological institute). red dots: core sites. k: kajak core. r: rumohr core. stippled lines: marine-terminating margins of flade isblink. b: bathymetry data points with colour code for 20 m depth intervals overlain on sar satellite image. depth data sources are from the 2015 asp field season (echo soundings and ctd based data) supplemented by data collected in 2006 by naja mikkelsen (geus), yngve kristoffersen (university of bergen, norway), and rené forsberg (dtu space, denmark). 69 in the wandel hav. this ice appears to be at least 6 years old and may date back to 2004 following the total breakup in 2003. a similar area of older ice patches is situated between the prinsesse thyra ø and prinsesse margrethe ø (fig. 2). in mid-august 2012, the fast-ice cover at the mouth of the independence fjord and danmark fjord broke up, creating a wide-open connection to the arctic ocean in the area north-west of prinsesse thyra ø. freezeup later in 2012 created the constellation of multi-year ice coverage persisting until today. only the outer rim of the old fast-ice cover may occasionally be eroded by loss of ice fragments, which drift away with the arctic pack-ice drift. the flade isblink ice cap is mainly drained by two outlets along its western margin. a comparison of earlier maps (higgins 1991) and 2015 satellite-derived data indicates that the outlet immediately north of station nord has retreated about 13 km southwards between 1991 and 2015. higgins (1991) estimated the average flow speed of the outlet glacier to be of a few hundred metres per year. satellite radar data from the last decade confirm this estimate, but also indicate that the outlet glaciers have been surging (joughin et al. 2010). field work we collected sediment cores from 11 april to 2 may 2015 in the ice-covered fjord area (figs 2, 3). sampling took place up to about 30 km from the station along several transects determined according to ice conditions and existing knowledge of bathymetry. we used two snow mobiles with sledges (3–4-person team) for transport. we selected sampling sites based on a geo-referenced high-resolution radar satellite image (sentinel-1 sar from 31.01.2015) revealing areas of first and multi-year ice as well as glacier-front positions and larger icebergs (fig. 2). whenever possible, we targeted sites with thin first-year ice (about 1.0–1.2 m) and avoided sites covered by thick (>3 m) multi-year ice, which was very arduous to drill through. at each core site, we first removed the snow cover, which was usually at least 1 m thick, and used a 9 inch ice-auger to drill two or more overlapping holes in the ice, sufficiently large for the coring devices (fig. 3). at most locations, we measured water depths with an echo sounder, but at oceanographic stations we used ctd data (conductivity, temperature and depth). we sampled sediment with a kajak sediment corer (25–75 cm long tube, 45 mm inner diameter), a rumohr lot corer (50–100 cm long tube, 75 mm inner diameter), and at a few stations, with a van veen grab sampler (top 10 cm surface sediment). we used a tripod with a top-mounted hydraulic winch (pot hauler) connected to a petrol-driven power pack to retrieve the c. 60 kg rumohr corer, whereas kajak cores were retrieved with a hand winch (fig. 3). drilling large holes through 1–3 m of sea ice was time-consuming, and we found that kajak coring was the most efficient approach. we collected duplicate or triplicate kajak cores at most coring sites and one kajak core from each site was subsampled at 1 cm intervals at the villum research station. sea-ice cores were collected using a kovacs ice corer system for biochemical and taxonomic studies of the sea-ice algal communities. bathymetrical data apart from a single study of short sediment cores and water-depth measurements south-west and north-west of prinsesse dagmar ø close to station nord (nørgaardpedersen et al. 2008) very little is known about the bathymetry and sedimentation record of this remote area. bathymetrical datasets from the limited earlier field projects in the area have, for this study, been updated by new data from the 2015 asp field season. earlier data consist of echo soundings carried out in 2006 and reconnaissance data. the 2015 field season dataset consists of echo sounding data and ctd-derived bathymetry data. as the ctd data are corrected for water-column velocity differences (due to water masses with different salinities and temperatures), these may be considered the most accurate. water depths increase from c. 20 m near station nord to >150 m 20–30 km northwards (fig. 2b). in front of the glacier outlet margin 10–15 km north-east of station nord, a trough is found with depths up to 100–150 m right up to the glacier margin at 81°40´n. there is a shallow area between prinsesse dagmar ø and prinsesse thyra ø with water depths between 20 and 30 m, but the depth increases fig. 3. recovering a kajak core with seabed sediments. photograph: jesper hoffmann. 7070 to >100 m towards the north. south of prinsesse dagmar ø, a trough possibly connected to the mouth of the danmark fjord shows depths in the range of 130–150 m. the trough axis rises to about 115 m c. 10 km from the glacier outlet margin and reaches depths of maximum 50 m at the glacier margin. preliminary results and outlook a total of 37 sediment cores were retrieved from 17 sites along transects up to c. 30 km from station nord (fig. 2a). many of the kajak cores are only a few decimetres long. however, we recovered cores exceeding 0.5 m from more water-rich mud close to glacier margins and at the few rumohr core sites. a spatial study of sea-ice and productivity proxies including dinoflagellate cysts, diatoms, foraminifera, biomarker ip25 (a proxy for sea ice) and biogenic silica is currently being conducted for the 17 sampling sites, to establish a baseline of recent conditions that will serve as modern analogues for reconstructions of sea-ice variability and changes in oceanographic conditions during earlier time periods. selected sediment cores are being analysed for 210pb and 137cs content to estimate sedimentation rates at the study sites and establish a chronology for the topmost part of the cores. the first dating results show sedimentation rates in the order of 0.04–0.06 cm/y. preliminary investigations of the microfossil content of the sediments revealed calcareous benthic foraminifera and a few ostracods, particularly at the deeper sites, which allows for the possibility to use 14c dating to establish a robust chronology for the sediment records. furthermore, studies of the sites located north of station nord show a stronger marine influence, whereas the sites towards the north-west and south have a clear signal of freshwater or glacial influence. complementary to the climate proxy work, a characterisation of the protist communities is being undertaken by germination, growth tests, and molecular analyses (dna) targeting the two main groups of primary producers: diatoms and dinoflagellates. investigations of the sampled sea-ice cores showed that during the early part of the season’s field work the sea ice was barren of algae. this is attributed to light attenuation by the snow (average snow thickness of 1 m). in situ light measurements, and fluorescence measurements on 25 sea-ice core samples (bottom 5 cm) using a phytoplankton analyser showed no detectable photosynthetic activity (information from the phytobiology team, aarhus university). the field work provided us with some first insights into an ice-covered, very remote and large fjord system of eastern north greenland. the preliminary data confirm presence of biogenic remains and sedimentary signatures which can be used as proxies for palaeo-environmental reconstructions and deciphering of the younger part of the holocene climate history in this region. acknowledgements we thank kunuk lennert, jesper hoffman, egon frandsen and the late john lau for logistic support. we thank leif toudal (dmi) for making geo-referenced satellite sar images of the study area available. asp 2015 oceanographers igor dmitrenko and sergei kirilov are also acknowledged for sharing ctd depth data. we thank the station nord military personnel for their great hospitality and help. the field work was partly financed by the villum foundation (grant no. vkr023454 to sofia ribeiro) and by the arctic research centre, aarhus university. references funder, s. 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(eds): quaternary glaciations – extent and chronology – a closer look. developments in quaternary sciences 15, 699–714. amsterdam: elsevier. higgins, a.k. 1991: north greenland glacier velocities and calf ice production. polarforschung 60, 1–23. jakobsson, m. et al. 2012: the international bathymetric chart of the arctic ocean (ibcao) version 3.0. geophysical research letters 39, l12609, http://dx.doi.org/10.1029/2012gl052219. joughin, i., smith, b.e., howat, i.m., scambos, t. & moon, t. 2010: greenland flow variability from ice-sheet-wide velocity mapping. journal of glaciology 56 (197), 415–430. nørgaard-pedersen, n., mikkelsen, n. & kristoffersen, y. 2008: late glacial and holocene marine records from the independence fjord and wandel sea regions, north greenland. polar research 27, 209–221. perovich, d.k. & richter-menge, j.a. 2015: regional variability in sea ice melt in a changing arctic. philosophical transactions of the royal society a373, 2045, http://dx.doi.org/10.1098/rsta.2014.0165. rasmussen, l. 2004: set fra oven: et hjørne af grønland. vejret 100, 17–20. wadhams, p. 1986: the ice cover. in: hurdle, b.g. (ed.): the nordic seas, 21–78. new york: springer-verlag. authors’ addresses n.n.-p., s.r., n.m. & a.l., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: nnp@geus.dk m.-s. s., department of geoscience, aarhus university, høegh-guldbergs gade 2, dk-8000 aarhus c, denmark. geological survey of denmark and greenland bulletin 7, 2004, p 61-64 61 during the past decade the geological survey of denmark and greenland (geus) has carried out two major resource evaluations in the precambrian basement terranes of south and west greenland in order to locate potential areas of mineral deposits (steenfelt et al. 2000, 2004; stendal & schønwandt 2003; stendal et al. 2004). based on geological field work and geochemical and geophysical data, these evaluations have assessed the interplay between the magmatic, tectonic and metamorphic evolution in the study areas and their mineralising events. as a result of the second of these evaluations it is now possible to outline a succession of mineralising events in the northern part of the nagssugtoqidian orogen and in the disko bugt area of central west greenland (fig. 1), and relate them to the general archaean and palaeoproterozoic geological evolution of this region. however, uncertainties still exist concerning the age and detailed setting of many epigenetic mineralisations. geological background the precambrian rocks of central west greenland north of the north atlantic craton, consist of archaean orthogneisses and supracrustal rocks together with localised belts of juvenile palaeoproterozoic intrusive complexes and supracrustal rocks. the whole region was reworked to varying degrees during a major palaeoproterozoic orogenic event. in west greenland the palaeoproterozoic has traditionally been divided into the nagssugtoqidian orogen between kangerlussuaq and disko bugt (fig. 1) and the rinkian fold belt farther north. however, more recent studies suggest that these two belts are largely contemporaneous and probably precambrian mineralising events in central west greenland (66°–70°15´n) henrik stendal and adam a. garde geological survey of denmark and greenland bulletin 7, 61–64 (2005) © geus, 2005 ataneq ilulissat kan ge rlu ssu aq 70°n 68°n attu rifkol kangaatsiaq aasiaat qeqertarsuaq disko disko bugt itilliarsuk nuussuaq saqqaq eqi ataa arveprinsen ejland qasigiannguit naternaq nordre isortoq sisimiut sukkertoppen iskappe sarfartoq 53°w 51°w inland ice nordre s trøm fjor d nassuttooq arfers iorfik greenland fig. 1. simplified geological map of central west greenland with mineral occurrences (modified from van gool et al. 2002b). precious metals base metals iron and iron alloys speciality metals light metals industrial minerals gemstones town kangerlussuaq airport quaternary ice palaeogene basalt cambrian sarfartoq carbonatite palaeoproterozoic sisimiut charnockitte palaeoproterozoic arfersiorfik quartz diorite palaeoproterozoic supracrustal rocks archaean granite archaean undifferentiated rocks archaean supracrustal rocks cretaceous–palaeogene sediments represent the southern and northern parts of a single, e–wtrending, c. 1850 ma collisional orogen. geological descriptions and additional references may be found in garde & steenfelt (1999), connelly et al. (2000, in press), garde et al. (2002, 2004) and van gool et al. (2002a). the nagssugtoqidian orogen is dominated by reworked archaean orthogneisses in the south (between c. 66°30´n and 69°n; fig. 1) with original emplacement ages of around 2870–2700 ma (connelly et al. 2000). it also incorporates two juvenile, calc-alkaline palaeoproterozoic plutonic magmatic complexes, namely the 1920–1870 ma arfersiorfik quartz diorite and the more or less contemporaneous sisimiut charnockite (van gool et al. 2002a). these must have been emplaced during subduction and prior to the continent–continent collision, and the presence of a suture in the central part of the nagssugtoqidian orogen has therefore been suggested. based on new zircon and titanite u-pb age determinations of the archaean basement and mylonites in the disko bugt area, connelly et al. (in press) have recently proposed that the main suture between the rinkian and nagssugtoqidian components lies in the disko bugt area; in the new model these components would represent the northern and southern colliding continents. the southern part of the nagssugtoqidian orogen hosts a rift-related mafic dyke swarm (the kangâmiut dykes) dated at c. 2040 ma (van gool et al. 2002a); the central and northern parts of the orogen contain palaeoproterozoic metasedimentary and metavolcanic belts of limited geographical extent, including the naternaq supracrustal belt (østergaard et al. 2002) which has been dated at around 1950 ma. these belts are thought to have been deposited along the margin of the rifted archaean continent. the northern part of the nagssugtoqidian orogen, the area east of disko bugt, and the precambrian rocks of eastern nuussuaq are all characterised by very variable palaeoproterozoic reworking. the archaean gneisses in southern nuussuaq are about 3000 ma old and include the itilli diorite dated at 3030 ma (connelly et al. in press). the northern part of the area east of disko bugt hosts two archaean, e–w-trending, amphibolite facies metavolcanic and metasedimentary belts, including a large sill complex and a low-grade palaeoproterozoic sedimentary succession (fig. 1; garde & steenfelt 1999; marshall & schønwandt 1999; stendal et al. 1999). the northern of the two archaean belts was probably deposited unconformably on the gneisses of southern nuussuaq and has been dated at c. 2950 ma (garde & steenfelt 1999; connelly et al. in press). following the palaeoproterozoic orogeny, a suite of e–wtrending ultramafic lamprophyres and small plugs was emplaced in the ataa area about 1750 ma ago (larsen & rex 1992). much later, at c. 600 ma, the sarfartoq carbonatite complex and numerous lamprophyric and carbonatitic dykes were intruded at the southern margin of the nagssugtoqidian orogen (see jensen et al. 2004), an event that has been associated with the opening of the iapetus ocean (larsen & rex 1992). 62 age (ma) ~ 600 1775–1600 1870–1775 1920–1870 2000–1920 ~ 2040 ~ 2650 ~ 2800 ~ 3000 geological setting rifting – carbonatite rifting – kimberlite cooling of the crust from c. 600–400ºc main nagssugtoqidian metamorphism and de formation. pegmatite formation ~ 1800 ma subduction – calcalkaline magmatism. peak metamorphism in disko bugt region drifting and sedimentation rifting granite intrusion into gneiss continental rift or active continental margin continental rift or active continental margin table 1. proterozoic mineralising events and commodities in west greenland locality and type of mineralising event sarfartoq – igneous kangerlussuaq – igneous hydrothermal activity in the whole region from 66º to 70º15´n. albitisation in disko bugt area remobilisation of metals during metamorphism and deformation in the whole region. monazite and allanite in pegmatite arfersiorfik diorite and sisimiut charnockite. hydrothermal activity n the disko bugt region naternaq, nordre strømfjord supracrustal rocks (ataneq). exhalative volcanic massive sulphides kangâmiut dykes rifkol, at attu intrusive related mineralisation eqi and arveprinsen ejland. exhalative volcanic massive sulphides and hydrothermal activity southern part of nuussuaq (saqqaq) and itilliarsuk. exhalative volcanic massive sulphides and hydrothermal activity commodity niobium, tantalum, phosphor diamond mainly minor sulphide showings with minor base metal contents and low in gold copper-zinc, graphite and formation of magnetite during metamorphism in mafic rocks. cesium in pegmatite minor oxide occurrences (fe-ti-v) in arfersiorfik. in disko bugt region minor copper, zinc and gold occurrences banded iron formation, cu-zn and gold copper gold copper, zinc, lead, gold and banded iron formation copper, zinc, gold and banded iron formation timing of mineralising events the timing of mineralising events in relation to the geological evolution of the study area is summarised below and in table 1. > 3000 ma the oldest known mineral occurrence of the region, at itilliarsuk in southern nuussuaq, comprises syngenetic banded iron formation (bif) and associated semi-massive sulphide occurrences, whereas the age of epigenetic gold occurrences in the same area is not known (stendal et al. 2004). 2800–2000 ma a pb isotope study of pyrite in the disko bugt area has yielded an age of c. 2800 ma (stendal 1998). gold occurrences related to syngenetic, semi-massive sulphide occurrences at eqi are presumed also to be of this age. a gold mineralisation at attu is interpreted to be of middle archaean age. the age of the pb source in the associated ore minerals is compatible with that of the nearby rifkol granite (kalsbeek et al. 1984), suggesting gold mineralisation at around 2650 ma. the pre-nagssugtoqidian continental break-up is represented by the well-known c. 2040 ma kangâmiut dyke swarm. the dykes themselves carry copper mineralisation with no economic significance. 2000–1920 ma several 1950–1920 ma old supracrustal suites in the northern nagssugtoqidian orogen (van gool et al. 2002a), e.g. at naternaq and ataneq, contain syngenetic massive sulphide occurrences (fig. 2), and the latter area also contains prominent graphite deposits. unpublished pb isotopic compositions of magnetite from amphibolite and banded iron formation yield intercept ages around c. 1940 and 2140 ma, in agreement with the syngenetic nature of the sulphide occurrences. 1920–1870 ma the 1920–1870 ma arfersiorfik and sisimiut intrusive suites are associated with nagssugtoqidian convergence and subduction of oceanic lithosphere prior to the main collision. the only known mineral occurrences resulting from this major event are minor igneous oxides (magnetite, ilmenite), although remobilisation and recrystallisation of previously formed mineral occurrences may also have taken place. in the disko bugt region, epigenetic fault zones and shear zones with small occurrences of base metals, nickel and gold are hosted in palaeoproterozoic supracrustal rocks, and have yielded a pyrite pb-pb model age of c. 1900 ma (stendal 1998). this important mineralising event was probably contemporaneous with the palaeoproterozoic thermal event recorded by k-ar and ar-ar ages in parts of the disko bugt region (rasmussen & holm 1999). 1870–1600 ma the nagssugtoqidian collisional orogeny is dated by early fabric-forming deformation at 1860–1840 ma, large-scale folding at c. 1825 ma, and formation of steep shear belts in the nassuttooq region at c. 1775 ma (van gool et al. 2002a). unpublished pb-pb ratios obtained by stepwise leaching of allanite and monazite indicate emplacement of pegmatites at c. 1800 ma in both central (nassuttooq) and northern (attu and qasigiannguit) regions of the orogen. metamorphic conditions in the central part of the nagssugtoqidian orogen reached temperatures above 650°c at 1800 ma, declining to 540°c at c. 1740 ma and 420°c at c. 1670 ma (connelly et al. 2000; willigers et al. 2001). pbpb isochron ages of c. 1740 ma from magnetite in amphibolites probably represent closure of the u-pb system in magnetite, and post-date the formation of prominent shear belts at around 1775 ma. the magnetite was probably formed during several different stages of metamorphism. signs of hydrothermal activity have been observed along neto nne-striking fault zones throughout the region. the cooling history of the orogen indicates that hydrothermal activity in the temperature range 650°–400°c extended from the time of pegmatite emplacement at c. 1800 ma until c. 1600 ma. 63 fig. 2. outcrop of volcanogenic-exhalative, semi-massive sulphides of palaeoproterozoic age at naternaq in the northern nagssugtoqidian orogen, with formation of gossan and sulphide weathering. in the disko bugt region, lamproites and ultramafic lamprophyres intruded the basement gneisses in the ataa area at around 1750 ma, and extensive post-kinematic albitisation was reported by kalsbeek & taylor (1999). however, mineralisations of this age are not known. 600 ma the formation of diamondiferous kimberlites and related carbonatite nb-ta mineralisation took place around 600 ma in the sarfartoq area (jensen & secher 2004; jensen et al. 2004). concluding remarks mineralising events in the precambrian basement terrain of central west greenland are confined to certain periods, of which some are more economically promising than others. the most prospective areas are the archaean greenstone belts in the disko bugt region, especially with respect to synand epigenetic gold occurrences. known palaeoproterozoic metal occurrences are all small, and probably without economic significance, with the possible exception of graphite deposits in the palaeoproterozoic supracrustal rocks in the sisimiut – nordre strømfjord region. niobium-tantalum and diamond occurrences in the neoproterozoic, c. 600 ma alkaline intrusions are currently under investigation and may be economically feasible. references connelly, j.n., van gool, j.a.m. & mengel, f.c. 2000: temporal evolution of a deeply eroded orogen: the nagssugtoqidian orogen, west greenland. canadian journal of earth sciences 37, 1121–1142. connelly, j.n., thrane k., krawiec, a.w. & garde, a.a. in press: linking the palaeoproterozoic nagssugtoqidian and rinkian orogens through the disko bugt region of west greenland. journal of the geological society (london). garde, a.a. & steenfelt, a. 1999: precambrian geology of nuussuaq and the area north-east of disko bugt, west greenland. geology of greenland survey bulletin 181, 6–40. garde, a.a., connelly, j.n., krawiec, a.w., piazolo, s. & thrane, k. 2002: a coastal survey in the southern part of the palaeoproterozoic rinkian fold belt, central west greenland. geology of greenland survey bulletin 191, 33–38. garde, a.a., connelly, j.n., grocott, j., hand, m., mccaffrey, k.j.w. & thrane, k. 2004: crustal shortening and granite emplacement in the rinkian fold belt, west greenland, and implications for palaeoproterozoic laurentian evolution. danmarks og grønlands geologiske undersøgelse rapport 2004/17, 16–18. jensen, s.m. & secher, k. 2004: investigating the diamond potential of southern west greenland. geological survey of denmark and greenland bulletin 4, 69–72. jensen, s.m., secher, k., rasmussen, t.m. & schjøth, f. 2004: diamond exploration data from west greenland: 2004 update and revision. danmarks og grønlands geologiske undersøgelse rapport 2004/117, 90 pp. kalsbeek, f. & taylor, p.n. 1999: review of isotope data for precambrian rocks from the disko bugt region, west greenland. geology of greenland survey bulletin 181, 41–47. kalsbeek, f., taylor, p.n. & henriksen, n. 1984: ages of rocks, structures, and metamorphism in the nagssugtoqidian mobile belt in west greenland – field and pb-isotopic evidence. canadian journal of earth sciences 21, 1126–1131. larsen, l.m. & rex, d.c. 1992: a review of the 2500 ma span of alkalineultramafic, potassic and carbonatitic magmatism in west greenland. lithos 28, 367–402. marshall, b. & schønwandt, h.k. 1999: an archaean sill complex and associated supracrustal rocks, arveprinsen ejland, north-east disko bugt, west greenland. geology of greenland survey bulletin 181, 87–102. østergaard, c., garde, a.a., nygaard, j., blomsterberg, j., nielsen, b.m., stendal, h. & thomas, c.w. 2002: the precambrian supracrustal rocks in the naternaq (lersletten) and ikamiut areas, central west greenland. geology of greenland survey bulletin 191, 24–32. rasmussen, h. & holm, p.m. 1999: proterozoic thermal activity in the archaean basement of the disko bugt region and eastern nuussuaq, west greenland: evidence from k-ar and 40ar-39ar mineral age investigations. geology of greenland survey bulletin 181, 55–64. steenfelt, a., nielsen, t.f.d. & stendal, h. 2000: mineral resource potential of south greenland: review of new digital data sets. danmarks og grønlands geologiske undersøgelse rapport 2000/50, 47 pp. steenfelt, a., stendal, h., nielsen, b.m. & rasmussen, t.m. 2004: gold in central west greenland – known and prospective occurrences. geological survey of denmark and greenland bulletin 4, 65–68. stendal, h. 1998: contrasting pb isotopes of archaean and palaeoproterozoic sulphide mineralisation, disko bugt, central west greenland. mineralium deposita 33, 255–265. stendal, h. & schønwandt, h.k. 2003: precambrian supracrustal rocks and mineral occurrences, northeast disko bugt. danmarks og grønlands geologiske undersøgelse rapport 2003/24, 57 pp. stendal, h., knudsen, c., marker, m. & thomassen, b. 1999: gold mineralisation at eqi, north-east disko bugt, west greenland. geology of greenland survey bulletin 181, 129–140. stendal, h., nielsen, b.m., secher, k. & steenfelt, a. 2004: mineral resources of the precambrian shield of central west greenland (66° to 70°15´n). part 2. mineral occurrences. danmarks og grønlands geologiske undersøgelse rapport 2004/20, 212 pp. van gool, j.a.m., connelly, j.n., marker, m. & mengel, f. 2002a: the nagssugtoqidian orogen of west greenland: tectonic evolution and regional correlations from a west greenland perspective. canadian journal of earth sciences 39, 665–686. van gool, j.a.m. et al. 2002b: precambrian geology of the northern nagssugtoqidian orogen, west greenland. geology of greenland survey bulletin 191, 13–23. willigers, b.j.a., krogstad, e.j. & wijbrans, j.r. 2001: comparison of thermochronometers in a slowly cooled granulite terrain: nagssugtoqidian orogen, west greenland. journal of petrology 42, 1729–1749. 64 authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: hst@geus.dk geological survey of denmark and greenland bulletin 26, 2012, 33-36 33 cliff collapse at stevns klint, south-east denmark stig a. schack pedersen and tove damholt the scenic coastal cliff of stevns klint is a classical study locality that stretches for 15 km along the east coast of sjælland and it holds arguably the best exposed cretaceous−tertiary boundary in the world (fig. 1; damholt & surlyk 2012). the famous boundary separates the soft cretaceous chalk from the harder overlying tertiary bryozoan limestone (fig. 2), and the difference between the two rock types controls the character of the frequent cliff falls. the relatively soft chalk at the base is eroded by storm waves and is subject to general debris shedding. the overlying bryozoan limestone, with its hardgrounds and flint layers, is more resistant to erosion and is strong enough to form overhanging projections of the coastal cliff that result in large and small recurring collapses. the position of the cretaceous−tertiary boundary varies in altitude along the cliff from about 5 m below sea level in the southern part of the cliff to c. 35 m above sea level in the northern part. hence the southern part of the cliff mainly consists of hard bryozoan limestone whereas the central and northern parts consist of soft chalk overlain by bryozoan limestone. throughout the entire length of the cliff the carbonate rock is overlain by a few metres of glacial till. stevns klint has recently been proposed for inclusion on the world heritage site list (damholt & surlyk 2012) and as part of the nomination process a risk assessment of the frequency of cliff collapse was conducted. this paper describes the analysis of erosion and rockfall that formed the background of the risk assessment. cliff-collapse analysis of stevns klint the evaluation of the cliff-collapse risk at stevns klint included an analysis of the size and character of the present overhang, the vulnerability of the cliff to erosion and rockfall dimensions. the analysis was based on a photogrammetric investigation using a series of oblique photographs taken in april 2011. the cliff section was mapped in segments and detailed photogrammetric measurements were made using the software socketset in the photogrammetric laboratory at the geological survey of denmark and greenland. the result was stored in a gis database using the arcgis format (pedersen & strunck 2011). results from a previous photogrammetric investigation using oblique photographs taken in 1992 (surlyk et al. 2006) were compared with our results to describe the changes in the cliff profile over the past 20 years. the general conditions of cliff erosion are described and the rockfall dynamics for the various types of cliff collapse are characterised. karise hårlev hellested 1 2 3 denmark højerup gamle kirke stevns fyr tommestrup harvig storedal lilledal korsnæb kirkevig holtug boesdal stevnsfort 2 km 50 km 12°25´e 56°10´n knøsen sjælland fig. 1. map of the stevns klint region showing place names mentioned in the text. the inset map shows the distribution of localities with major landslides in denmark. red trinangles: the most hazardous slides. yellow triangles: clayey landslides. 1: lønstrup klint, 2: stevns klint, 3: møns klint. till limestone chalk clay sea beach talus overhang 10 m fig. 2. block diagram illustrating the morphology and rock types present in the stevns klint exposures. the succession consists of maastrichtian chalkish clay, with the famous cretaceous–tertiary boundary at the base, danian bryozoan limestone and weichselian glacial till. © 2012 geus. geological survey of denmark and greenland bulletin 26, 33–36. open access: www.geus.dk/publications/bull 3434 general erosion conditions at stevns klint the rate of erosion is an essential factor for the assessment of cliff collapse. we estimated the average rate of coastal erosion at stevns klint to be 15 cm/year, based on comparisons of the position of the coastline in 1891 with that in 2010, provided by the national survey and cadastre. the highest rate (c. 35 cm/year) was found in the northern part of the cliff between storedal and lilledal (fig. 1). in some places, progradation of the coastline occurs where beach ridges form or where old cliff slides act as wave breakers. beach ridges are especially found at korsnæb in the south (fig. 1), where the accumulation rate amounts to 12 cm/year. a landslide north of kirkevig and south of storedal (fig. 1) has resulted in a coastline progradation of 7–9 cm/year. limestone quarrying over the past few hundred years has also altered the cliff profile. numerous small quarries were established directly on the cliff with quarry walls inclined towards the base of the limestone and resulted in considerable overhang. the quarries can contribute to cliff points projecting above the beach. the erosion rate at stevns klint is modest when compared with the highest rate of coastal erosion in denmark, which is 1.25 m/year at lønstrup klint that is formed in quaternary sediments (fig. 1; pedersen 1986). types of cliff collapses at stevns klint the largest type of cliff collapse at stevns klint can be called cliff slides (fig. 3). this type of slide includes a major part of the cliff for several tens of metres along the coastline which slide simultaneously. in general a cliff slide takes place along a steeply dipping surface and it involves a volume of 5000 to 10 000 m3. prior to a cliff-slide a whole section of the chalk has been eroded away at sea level, creating an undercut below the limestone and resulting in a considerable overhang. the size of the overhang has to be around 10–12 m and the thickness of the limestone and till in the overhang around 15–18 m before it becomes unstable. the most famous event at stevns klint occurred in 1928, when a large cliff slide dramatically tore away the choir of fig. 3. cliff-collapse types at stevns klint. the most hazardous type is the cliff slide involving volumes of 5000–10 000 m3. the volume of a cliff fall amounts only to 500–1500 m3, but still a cliff fall is an impressive sight with the bedding completely preserved in a displaced cliff fragment. the rockfall is the general type of cliff collapse and is illustrated by two examples, of which the plate-like overhang corresponds to a recent cliff collapse at knøsen. the rock-bedding exfoliation generates platy 1–10 m3 limestone blocks, which drop from the underside of the overhang. the main erosion of the cliff is caused by debris shedding, which produces the talus at the foot of the cliff. cliff with projection overhang before collapse 40 m rockfall after collapse drop of cliff fragments due to exfoliation erosion of the roof in breakers caves rock-bedding exfoliation 25 m rockfall cliff with overhang before collapse 40 m cliff slide after collapse cliff slide debris shedding rockfall cliff fall cliff slide after collapse a projecting cliff above overhang before collapse collapse of thin overhang constituting the floor of an old limestone pit 30 m 40 m talus cones with accumulation of fragments due to seasonal exfoliation erosion 20 m rockfall after collapse 35 højerup gamle kirke (a medieval church) that rested on a 15 m thick overhang. the slide continued several tens of metres out into the sea (fig. 4). shortly after the incident the cliff was protected by wave breakers to inhibit wave erosion and to secure the rest of the old church. prior to the spectacular slide, the undercutting erosion had lasted for 160 years following an earlier slide in 1767 (rasmussen 1967). before the cliff collapse in 1928, the overhang amounted to c. 12 m measured perpendicularly to the orientation of the cliff and a similar overhang size is estimated for the collapse in 1767. an overhang of c. 12 m is thus regarded as a critical size for cliff slides where the overhang thickness is c. 15 m. a significant cliff slide has occurred at tommestrup (figs 1, 5a) and another one in the inner part of harvig. at both sites cliff-slide deposits will protect the cliff from erosion for the next decades. in order of decreasing size, the next type of large collapse is the cliff fall (fig. 3). cliff falls only involve isolated projecting parts of the cliff with a considerable overhang and their volumes are 500–1500 m3. the cliff fall is a simple drop of part of the cliff where the bedding of the fallen block is preserved after the fall. the average relation between thickness and depth of the overhanging limestone prior to cliff fall is 3 to 2 as illustrated in the cliff fall below the lighthouse of stevns fyr (figs 1, 5b). the third type of collapse is classified as rockfalls (fig. 3), which in this analysis varies in size from 1 to 500 m3. this type may occur together with block slides, in which bigger blocks are displaced together with finer-grained material such as clayey till and sand. the resulting aggradation of material on the shore after a rockfall is a chaotic breccia. in general, rockfalls leave a concave escarpment in the cliff and the aggradation of debris results in a cone expanding out into the sea. platy rockfall, caused by rock-bedding exfoliation, involves rock volumes of around 1 m3 and is a significant type of collapse at stevns klint (fig. 3). rock-bedding exfoliation occurs in caves eroded by storm waves in the lower part of the cliff. layers are peeled off from the roofs of the cave due to exfoliation, and dish-shaped fragments of rock drop to the cave floor. caves are prominent north of boesdal (fig. 1) and fig. 4. aerial photograph of the cliff section at højerup gamle kirke (medieval church). fig. 5. cliff sections showing various types of rockfalls at stevns klint. a: cliff slide at tommestrup. b: cliff fall at the lighthouse of stevns fyr. a b 3636 north of stevnsfort where they are formed in limestone and extend up to 9 m into the cliff. caves formed in chalk occur south of storedal and at holtug. finally, fragments of rock may fall from the cliff due to general shedding of debris (fig. 3). in general, the material resulting from debris shedding is pebbles and cobbles, whereas clasts of boulder-size are loosened only occasionally from the cliff surface. the result is seen as an apron of talus along the foot of the cliff, typically with irregular conical shapes. debris shedding results from contraction and expansion due to seasonal variation of frost and thaw in winter time and desiccation during the summer. the rate of erosion due to debris shedding is equal to the average rate of erosion, i.e. 15 cm/year. assessment of cliff collapse hazards at stevns klint the size of the overhang is considered to be the most important factor in the assessment analysis of rockfall hazards at stevns klint. however, the likelihood of a collapse also depends on the thickness of the overhang. the shear strength of a rock increases with increasing normal pressure, which increases with the thickness of the rock. therefore an overhang with a thickness of 20 m is more stable than one with a thickness of 10 m. the degree of exposure to the sea is also important, primarily because the exposed parts of the cliff are affected by wave erosion and secondly because the softness of chalk increases by salt water spray (mortimore et al. 2004). finally, the assessment of risk is proportional to the number of visitors. a more detailed description of the assessment is provided in a full report (pedersen & strunck 2011). after our analysis, a rockfall of c. 300 m3 occurred at knøsen (fig. 1). the rockfall was discovered on 6 november 2011. it was caused by the collapse of a 5 m overhang with a thickness of 2 m that constituted the floor of an old limestone quarry (figs 3, 6). the floor carried the weight of removed overburden, in the form of clayey till and chertrich limestone rubble.  the collapse shows that an overhang with a thickness of just 2 m can become unstable even if the overhang is only 5 m. in other instances an overhang with a thickness of 15 m can be stable even with an undercut of more than 10 m. discussion and conclusions our cliff-collapse hazard assessment of stevns klint is based on an investigation of cliff collapse processes along the cliff. five types of cliff collapse are recognised: cliff slides, cliff fall, rockfall, platy rockfall caused by rock-bedding exfoliation and debris shedding. the cliff slides have large volumes and occur at low frequencies (1/100 years) whereas debris shedding of low volume occurs at high frequencies (seasonal). the resulting average rate of erosion is 15 cm/ year. rockfalls may be initiated by cycles of frost and thaw, saturation of cracks following rainy periods or desiccation after dry periods, but actual prediction of cliff collapses is not possible at present. references damholt, t. & surlyk, f. 2012: nomination of stevns klint for inclusion in the world heritage list, 159 pp. st. heddinge: østsjællands museum. mortimore, r.n., stone, k.j., lawrence, j. & duperret, a. 2004: chalk physical properties and cliff stability. in: mortimore, r.n. & duperret (eds): coastal chalk cliff instability. geological society (london), engineering geology special publications 20, 75–88. pedersen, s.a.s. 1986: rubjerg klint, rubjerg klit, årsag og virkning. varv 3, 84–98. pedersen, s.a.s. & strunck, m.n. 2011: vurdering af fjeldskredsrisiko på stevns klint. danmarks og grønlands geologiske undersøgelse rapport 2011/93, 100 pp. rasmussen, h.w. 1967: skrivekridtet og kalkstenene. in: nørvang, a. & meyer, t.j. (eds): danmarks natur 1, landskabets opståen, 131–160. copenhagen: politikens forlag. surlyk, f., damholt, t. & bjerager, m. 2006: stevns klint: uppermost maastrichtian chalk, cretaceous–tertiary boundary, and lower danian bryozoan mound complex. bulletin of the geological society of denmark 54, 1–48. fig. 6. the cliff collapse at knøsen, which occurred in november 2011. the weight of fill deposited on top of the section added to the load of the overhang. authors’ addresses s.a.s.p., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: sasp@geus.dk t.d., østsjællands museum, højerup bygade 38, dk-4660 st. heddinge, denmark. geological survey of denmark and greenland bulletin 42, 2018, 85-113 85 petroleum potential of the upper jurassic hareelv formation, jameson land, east greenland jørgen a. bojesen-koefoed, morten bjerager, h. peter nytoft, henrik i. petersen, stefan piasecki and anders pilgaard the marine, mudstone-dominated hareelv formation (upper jurassic) of jameson land, east greenland is a representative of the widespread kimmeridge clay formation equivalents, sensu lato, known from the greater north atlantic region, western siberia and basins off eastern canada. these deposits constitute the most important petroleum source-rock succession of the region. the present study reports petroleum geochemical data from the 233.8 m thick succession penetrated by the fully cored blokelv-1 borehole, and includes supplementary data from outcrop samples and other boreholes in jameson land. the succession consists of basinal mudstone intercalated with a significant proportion of gravity-flow sandstones, both in situ and remobilised as injectites. the mudstones are generally rich in organic carbon with values of toc reaching nearly 19 wt% and high pyrolysis yields reaching values of s2 up to nearly 43 kg hc/ton. hydrogen indices are up to 363. the data presented herein demonstrate that weathering of abundant pyritic sulfur adversely affects the petroleum potential of the kerogen in outcrop samples. the succession is thermally immature to early mature, except where intrusions have locally heated adjacent mudstones. the documentation of rich gas/oil-prone upper jurassic successions in jameson land is important for the assessment of the regional petroleum potential, including the north-east greenland continental shelf. keywords: east greenland, upper jurassic, petroleum source rock, generation potential, thermal maturity. ___________________________________________________________________________ j.a.b.k., m.b., h.p.n. & s.p., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: jbk@geus.dk h.i.p., total exploration & production denmark, amerika plads 29, 2100 copenhagen k, denmark s.p., also geological museum, natural history museum of denmark, university of copenhagen, øster voldgade 5-7, dk-1350 copenhagen k, denmark a.p., dansk miljørådgivning a/s, karolinevej 17, dk-4200 slagelse, denmark oxfordian–ryazanian marine shales form the most important petroleum source rocks of the prospective basins of the north atlantic region. these include the basins of the greater north sea area, the barents shelf, the basins west of ireland and the shetland islands, the jeanne d’arc and flemish pass basins off eastern canada as well as the basins of western siberia, and probably also the hitherto untested basins off east and north-east greenland. the deposits have been the subjects of numerous studies (e.g. von der dick et al. 1989; miller 1990; chakhmakhchev et al. 1994; klemme 1994; telnæs et al. 1994; fowler & mcalpine 1995; isaksen & ledje 2001; ineson et al. 2003; justwan & dahl 2005; justwan et al. 2005, 2006a,b; petersen et al. 2010), and age-equivalent rocks crop out onshore north-east greenland. despite their importance for future exploration in the east and north-east greenland offshore region, published in-depth studies of their nature and petroleum potential have hitherto been scarce, with a few notable exceptions (requejo et al. 1989; christiansen et al. 1992; strogen et al. 2005). during the © geus, 2018. geological survey of denmark and greenland bulletin 42, 85–113. available at: www.geus.dk/bulletin42 mailto:jbk@geus.dk http://www.geus.dk/bulletin42 8686 period 2008–2010, the geological survey of denmark and greenland (geus) drilled three fully cored boreholes to depths of more than 200 m targeting the upper jurassic – lower cretaceous shale succession in east and north-east greenland. the deposits are roughly timeequivalents of the kimmeridge clay formation of the wessex basin, uk, and comparable core material exists from the kimmeridge clay formation close to its type section in dorset (morgans-bell et al. 2001). however, the east and north-east greenland cores offer a unique opportunity to study nearly the full succession, which ranges in age from oxfordian to ryazanian, in an area remote from other studied outcrops and wells. the first of these boreholes to be drilled was blokelv-1 in jameson land, which covers the succession from the oxfordian to the lower volgian. the objective of this paper is to present a comprehensive overview of the petroleum potential of the upper jurassic – lower cretaceous shale succession in jameson land, based on new evidence from the blokelv-1 core, while including other available data from both outcrops and earlier shallow, fully cored boreholes drilled in jameson land (fig. 1). details of the drilling, sedimentology and stratigraphy of the blokelv-1 borehole can be found in bjerager et al. (2018a, b, this volume) and alsen et al. (2018, this volume). geological setting the jameson land basin is interpreted as a rift basin bounded by faults against caledonian basement rocks to the west and against the uplifted basement rocks of liverpool land to the east (hamann et al. 2005, and references therein). the nature of the northern and southern limits of the basin is unclear, although deep-seated nw– se-trending faults on the northern part of jameson land fig. 1. geological map of jameson land showing outcrop sampling areas mentioned in the text, locations of the ‘old boreholes’ (lollandselv-1, -2, falsterelv-1, jyllandselv-1, sjællandselv-1, -2, -3) and the blokelv-1 borehole. inset shows the location of southern jameson land in east greenland. ; ; ; ; ; ; jyllandselv sjællandselvblokelv h es te elv lollandselv lollandselv-2 lollandselv-1 falsterelv-1 jyllandselv-1 blokelv-1 sjællandselv-1-2 -3 h urry in le t goniomyakløft katedralen fortet jameson land liverpool land fjern søer falsterelv major dyke/sill fault ice river (selected) quaternary hesteelv fm raukelv fm hareelv fm olympen fm fossilbjerget fm/pelion fm neill klinter gp kap stewart gp triassic basement 20 km ?devonian 70°30'n 71°n 22°w23°w24°w jbk 1 s c o r e s b y s u n d 87 have been interpreted to have influenced the depositional patterns in the jurassic and may mark the transition to the traill ø basin (dam et al. 1995). the tectonic style changes north of jameson land towards the wollaston forland basin where the structural style is dominated by westerly tilted fault blocks (vischer 1943; surlyk 1978). the fill of the jameson land basin consists of devonian to cretaceous sediments with a composite thickness of up to 17 km. the basin fill is gently tilted (2–3°) towards the sw, and younger deposits are thus primarily present in southern jameson land. the succession includes oxfordian–ryazanian marine organic-rich mudstones (christiansen et al. 1992, 1993), broadly age-equivalent to the widespread, rich petroleum source-rock successions of the offshore areas of northwest europe. such deposits are also widespread in other areas of east and north-east greenland, where they are referred to several lithostratigraphic units (surlyk 2003). in jameson land, this succession is referred to the olympen, hareelv, raukelv and hesteelv formations (fig. 2); bjerager et al. (2018c, this volume) provide details of the stratigraphy and regional geology. the upper part of the olympen formation is represented by shallow marine deltaic sandstones with a transition downslope to bioturbated silty mudstones and gravity-flow sandstones. in the basinal settings, the formation passes up into the lower slope to basinal hareelv formation (hareelv = hare river; lower oxfordian – lower volgian) which is up to 200 m thick in the eastern part of the basin and may reach thicknesses of 300–500 m more centrally in the basin. organic-rich marine mudstones are largely restricted to the katedralen member (oxfordian–kimmeridgian), fig. 2. stratigraphic framework (modified from bjerager et al. 2018c, this volume) showing the stratigraphic range of the cored blokelv-1 borehole. c re ta ce ou s lo w er u pp er ju ra ss ic o xf or di an l l l l m u u u u sy st em se rie s te th ys (s ta nd ar d) bo re al su bs ta ge m ki m m er id gia n be rr ias ian ti th on ian ry az an ian vo lgi an age (ma) 140 150 160 145 155 stage s njameson land hareelv fm? hareelv fm hades mb olympen fm zeus mbzeus mb sa katedralen mb sjællandselv mb shallow marine sandstone shelf transition silty–sandy mudstones and heteroliths deep marine sandstone deep shelf (basin) mudstone mass-flow sandstone remobilised/injected sandstone raukelv fmraukelv fm source rockprograding unit bl ok el v1 jbk fig 2 hiatus 8888 whereas the overlying sjællandselv member is dominated by coarser clastic facies; the salix dal member, which overlies the sjællandselv member, is dominated by silty mudstones. deposition of the katedralen member mudstones took place in a deep-water marine setting and was interrupted by frequent deposition of gravity-flow sandstones, commonly remobilised post-depositionally as injectites (surlyk & noe-nygaard 2001). the raukelv formation, which succeeds the hareelv formation, represents rapid coastal progradation of coarse clastic deposits into the deep basinal areas and is capped by a prominent erosional unconformity. the overlying hesteelv formation, exposed near the southern coast of jameson land, fills a c. 10 km wide, u-shaped channel with a southwards plunging axis. black mudstones of the crinoid bjerg member occupy the deepest part of the channel (surlyk et al. 1973; surlyk 1973). these deposits accumulated during regional drowning in the ryazanian. the crinoid member of the hesteelv formation is thus age-equivalent to the more organic-rich intervals of the north sea source-rock successions (ineson et al. 2003). samples and methods samples used for the present study include material from the fully cored blokelv-1 borehole, which was drilled in 2008, as well as samples from shallow fully cored boreholes drilled in 1983 and 1993, and samples collected from outcrops in jameson land over several decades for a variety of purposes (fig. 1, table 1). the stratigraphic range covered by all boreholes and their mutual stratigraphic relationships are shown in fig. 3, which also provides a general indication of the stratigraphic intervals covered by outcrop samples included in this study. not all outcrop samples have been dated, but a general indication of the age of samples collected from the various sampling areas is given in table 1. outcrop samples represent the hareelv formation (predominantly the katedralen member with a few samples from the salix dal member) and the hesteelv formation. for most samples, total carbon (tc, wt%), total organic carbon (toc, wt%) and total sulfur (ts, wt%) were determined by combustion in a leco cs-200 induction furnace. toc was determined after elimination of carbonate-bonded carbon by prolonged hcl treatment. petroleum potential was determined by rockeval-type pyrolysis using a source rock analyzer (sra) instrument, manufactured by humble instruments and services (presently weatherford), and calibrated against the ifp160000 standard, running a blank and an inhouse control standard for every 10 analyses. for some older samples, these parameters were determined using similar methods and comparable instruments such as the rock-eval 2 or rock-eval 6. based on the experience at the geus laboratory, analyses of samples from the same successions carried out over a period of more than 25 years show very good agreement, despite differences in c or e core-id drilled latitude longitude depth (m) blokelv-1 511101 2008 70°45.305´ 23°40.430´ 233.8 sjællandselv-1 303114 1983 70°42.0' 23°25.0' 30.53 sjællandselv-2 303115 1983 70°42.0' 23°25.0' 40.28 sjællandselv-3 303116 1983 70°42.0' 23°25.0' 44.62 lollandselv-1 303139 1993 70°55.63' 23°50.53' 101.2 lollandselv-2 303140 1993 70°58.00' 23°49.02' 85.5 falsterelv-1 303141 1993 70°52.91' 23°47.94' 101.2 jyllandselv-1 303142 1993 70°47.37' 23°39.13' 73.4 total o ut cr op blokelv area n.a. n.a. n.a. n.a. n.a. falsterelv area n.a. n.a. n.a. n.a. n.a. fortet area n.a. n.a. n.a. n.a. n.a. gonyomyakløft area n.a. n.a. n.a. n.a. n.a. jyllandselv area n.a. n.a. n.a. n.a. n.a. katedralen area n.a. n.a. n.a. n.a. n.a. lollandselv area n.a. n.a. n.a. n.a. n.a. unspecified area n.a. n.a. n.a. n.a. n.a. hesteelv area (hesteelv fm) n.a. n.a. n.a. n.a. n.a. total age samples analysed m. oxfordian – l. volgian 139 u. kimmeridgian – l. volgian 40 u. kimmeridgian – l. volgian 21 u. kimmeridgian – l. volgian 43 l. oxfordian 30 l. – m. oxfordian 38 l. – m. oxfordian 58 u. oxfordian 83 452 u. oxfordian – kimmeridgian 3 u. oxfordian – kimmeridgian 18 kimmeridgian–volgian 43 oxfordian 3 oxfordian–kimmeridgian 11 oxfordian 26 oxfordian 12 n.a. 4 ryazanian 8 128 jbk table 1 table 1. overview of cored boreholes and outcrop samples 89 instrumentation used for analysis, provided similar calibration standards are used. note that data published by requejo et al. (1989) have been omitted here since they seem to include a calibration error with respect to tmax. particulate blocks for reflected light microscopy were prepared according to international standards (taylor et al. 1998). vitrinite reflectance (ro) measurements (random) were made using monochromatic light (546 nm) after calibration with standards of 0.515 %ro and 0.92 %ro. ro for all samples was calculated from a population of vitrinite reflectance recordings selected from the measured ro histogram. a number of samples were also qualitatively inspected in reflected white light and fluorescence-inducing blue light using a zeiss microscope. solvent extraction of powdered samples (<250 µm) was carried out with methanol/dichloromethane 7:93 vol./ vol. as solvent using a soxtec equipment. asphaltenes were precipitated with a 40-fold excess of n-pentane. maltene fractions were separated into saturated, aromatic and nso fractions by medium-pressure liquid chromatography (mplc) following a procedure inspired by radke et al. (1980). gas chromatography of saturated extract fractions was carried out using a shimadzu gc-2010 instrument with splitless injection and a zb-1 capillary column (25 m × 0.25 mm inner diameter, film thickness 0.10 μm). the temperature programme was 5°c/min. from 80° to 300°c, followed by 15 min. at 300°c. gas chromatography – mass spectrometry was carried out using an agilent 6890n gas chromatograph connected to a waters (micromass) quattro micro gc tandem quadropole mass spectrometer. a phenomenex zb-5 column (30 m × 0.25 mm inner diameter, film thickness 0.10 μm) was used. the injection temperature c re ta ce ou s lo w er u pp er ju ra ss ic o xf or di an l l l l m u u u u sy st em se rie s te th ys (s ta nd ar d) bo re al su bs ta ge m ki m m er id gia n be rr ias ian ti th on ian ry az an ian vo lgi an age (ma) 140 150 160 145 h es te el v fo rt et outcropscored borehole sections ka te dr ale n lo lla nd se lv bl ok el v, fa lst er el v jy lla nd se lv 155 stage bl ok el v1 h a fm h a fm h a fm h a fm h a fm h e fm h a fm o l f m o l f m o l f m o l f m h a fm lo lla nd se lv1 (3 03 13 9) lo lla nd se lv2 (3 03 14 0) fa lst er el v1 (3 03 14 1) jy lla nd se lv1 (3 03 14 2) sjæ lla nd se lv1, -2 , 3 (3 03 11 4, 30 31 15 , 3 03 11 6 pa rt ial ly ov er lap pi ng ) jbk 3 fig. 3. stratigraphic relationships and intervals sampled by cores and outcrops. ha fm: hareelv formation. he fm: hesteelv formation. ol fm: olympen formation. the greenland geological survey (ggu) registration number is indicated for the older suite of boreholes. 9090 was 70°c (2 min. hold). the temperature programme was 30°c/min. from 70° to 100°c and 4°c/min. from 100° to 308°c followed by 8 min. at 308°c. argon was used as the collision gas in ms-ms runs. the saturated hydrocarbons were analysed by gc–ms in sim-mode and by gc–ms–ms using relevant parent–daughter transitions for c26–30 steranes and c27–35 hopanes. four separate gc–ms and gc–ms–ms methods were applied to optimise data quality. samples were dissolved in isooctane (3 mg / ml). for a number of samples, ‘vintage’ biological marker data exist, but these have been omitted in order to ensure that data can be compared. hence only biological marker data produced according to the above protocol, or similar, have been included. results outcrop samples the hareelv formation is commonly poorly exposed (fig. 4), but outcrops were sampled at a number of different locations, several situated close to borehole drill sites (fig. 1, table 1); samples were primarily collected from the older, oxfordian–kimmeridgian, parts of the succession. in general, the samples are rich in organic carbon, showing average toc values from about 2 wt% to well over 8 wt% (table 2). total sulfur data are only available from the fortet area, where the deposits show values in the range 0.3 to 3.1 wt% total sulfur (ts). the petroleum potential as indicated by the pyrolysis s2-peak fig. 4. air photo (helicopter) of the central part of jameson land, viewed towards the east, showing the characteristic plateau-like landscape featuring flat-topped hills formed by the hareelv formation. the recessive nature of the hareelv formation serves to make good outcrops relatively scarce. the snow-clad mountains of liverpool land are visible in the distance. 91 shows wide variation with average values ranging from less than 1 kg/ton to nearly 12 kg/ton, giving average hydrogen index (hi) values in the range from 24 to 189 (fig. 5, table 2), with only a few individual samples exceeding 200. tmax and the production index (pi) consistently indicate pre-oil window maturity, except for a few samples from the fortet and katedralen outcrops that have been locally heated by intrusions, showing deviating data due to elevated maturity and oil staining (fig. 5, table 2). an attempt to establish the amount of nonreactive carbon and the average hydrogen index of the reactive fraction by cross-plotting s2 versus toc (dahl et al. 2004) using linear axes was unsuccessful since wide scatter precluded the construction of reliable trendlines. systematic palynofacies data are not available, but from the study of slides for palynological dating it is known that the deposits are often very rich in woody terrestrial organic matter. in summary, despite high levels of toc, outcrop samples in general show rather poor petroleum potential. the majority of the samples may be classified as primarily gas-prone, containing predominantly type iii kerogen, or even as non-source deposits containing predominantly inert kerogen. a few samples show hi values between 200 and 300 and may be classified as gas/ oil-prone. the hesteelv formation, stratigraphically overlying the hareelv formation and age-equivalent to the more organic-rich upper interval of the north sea source-rock succession, is represented by a total of eight samples (table 1). the organic carbon content is close to 2 wt%, s2 ranges from 2 to 5 kg/ton, giving hydrogen index values in the range 118–225 with an average of 149. all samples are thermally immature with tmax close to 430°c and pi close to 0.05 (fig 5, table 2). ’old boreholes’ the sjællandselv-1, -2, -3, lollandselv-1, -2, falsterelv-1 and jyllandselv-1 fully cored boreholes resulted from drilling campaigns in the 1980s and 1990s and predate the drilling of the blokelv-1 borehole in 2008; they are referred to here as ‘old boreholes’. detailed correlation of these cored sections can be found in bjerager et al. 2018c, this volume. lollandselv-1 (core 303139), lollandselv-2 (core 303140), falsterelv-1 (core 303141). the three boreholes represent the lower to lowermost upper oxfordian succession, mainly referred to the olympen formation but with the transition into the hareelv formation recorded in the top part of the fasterelv-1 borehole (fig. 3, table 1); in combination, they provide a fairly complete representation of this interval. the succession is dominated by sandy deposits with subordinate mudstones that are typically silty, poorly laminated or structureless. in general, the samples collected from the cores are rich in organic carbon, showing average values between 3.8 wt% and 4.8 wt% (table 3). the average petroleum potential, as indicated by s2, varies from 1.8 kg/ton to 4.8 kg/ton, with hydrogen index (hi) values ranging from 48 to 92 (fig. 6, table 3). tmax and the production index (pi) values indicate pre-oil window maturity. jyllandselv-1 (core 303142). this borehole penetrated the uppermost upper oxfordian – lower kimmeridgian part of the succession (fig. 3, table 1). the succession is dominated by mudstones, parts of which are laminated, with subordinate sandstones. the samples are rich in organic carbon with an average of 7.6 wt% (table 3). the petroleum potential, as indicated by s2, averages 15.6 kg/ ton with an average hi value of 196 (fig. 6, table 3). tmax and the production index (pi) generally indicate pre-oil min max mean min max mean min max mean min max mean min max mean min max mean hi pi no. (samples) toc tmax s1 s2 bl ok el v ar ea fa lst er el v ar ea fo rt et a re a g on io m ya kl øf t a re a jy lla nd se lv ar ea ka te dr ale n ar ea lo lla nd se lv ar ea 3 18 43 3 11 26 12 4.07 2.35 0.60 5.38 3.31 2.36 2.14 5.01 7.85 8.79 11.1 10.97 11.50 7.45 4.57 4.27 4.26 8.17 7.19 7.30 4.54 429 422 310 420 419 421 423 432 433 609 423 430 441 450 431 427 436 421 424 429 434 0.61 0.02 0.00 0.14 0.16 0.02 0.01 0.64 0.30 2.33 0.75 0.46 1.62 0.04 0.63 0.09 0.49 0.47 0.35 0.42 0.03 8.04 0.76 0.03 1.78 3.33 0.50 0.40 9.01 8.20 13.94 17.35 17.84 18.25 1.65 8.58 3.22 3.57 11.49 11.27 4.80 0.99 173 28 2 33 76 15 13 198 180 234 191 231 171 43 189 75 88 127 150 59 24 0.07 0.02 0.00 0.03 0.02 0.03 0.01 0.07 0.05 0.53 0.07 0.06 0.27 0.05 0.07 0.03 0.15 0.05 0.04 0.07 0.03 u ns pe cif ie d ar ea s h es te el v ar ea (h es te el v fm ) 4 8 1.44 1.25 3.19 2.32 2.09 1.96 432 427 434 436 433 430 0.02 0.05 0.07 0.22 0.04 0.14 0.34 0.66 0.81 4.94 0.58 2.46 18 53 36 225 28 120 0.04 0.04 0.08 0.07 0.06 0.06 jbk table 2 table 2. summary of petroleum potential data for outcrop samples, all data 9292 excellent ex ce lle nt good g oo d fa ir po or poor blokelv falsterelv fortet goniomyakløft jyllandselv katedralen lollandselv unnamed location hesteelv a staining or contamination low maturity conversion c immature oil condensate wet gas dry gas 0.1 1 10 100 toc (%) 0.1 1 10 100 400 450 500 550 tmax (°c) 0 100 200 300 400 500 600 700 800 900 1000 h yd ro ge n in de x s2 (k g hy dr oc ar bo ns /to n ro ck ) pr od uc tio n in de x (s 1/ (s 1+ s2 )) 380 390 400 410 420 430 440 450 460 470 480 490 500 510 520 tmax (°c) 0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1 b jbk 5 fig. 5. organic geochemical screening data for outcrop samples. a: tmax vs. hydrogen index. b: toc versus s2. c: tmax vs. production index. 93 fig. 6. organic geochemical screening data for ‘old borehole’ samples. a: tmax vs. hydrogen index. b: toc vs. s2. c: tmax vs. production index. sjællandselv-1 (303114) sjællandselv-2 (303115) sjællandselv-3 (303116) lollandselv-1 (303139) lollandselv-2 (303140) falsterelv-1 (303141) jyllandselv-1 (303142) immature oil condensate wet gas dry gas staining or contamination low maturity conversion 0.1 1 10 100 toc (%) 0.1 1 10 100 excellent ex ce lle nt good g oo d fa ir po or poor 400 450 500 550 tmax (°c) 0 100 200 300 400 500 600 700 800 900 1000 h yd ro ge n in de x s2 (k g hy dr oc ar bo ns /to n ro ck ) pr od uc tio n in de x (s 1/ (s 1+ s2 )) 380 390 400 410 420 430 440 450 460 470 480 490 500 510 520 tmax (°c) 0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1 a b c jbk 6 9494 window maturity, although the maturity level seems significantly higher than that recorded in the lollandselv-1, lollandselv-2 and falsterelv-1 successions. sjællandselv-1 (core 303114), sjællandselv-2 (core 303115), sjællandselv-3 (core 303116). these three boreholes were drilled close together and overlap stratigraphically, together covering the upper kimmeridgian – lowermost volgian part of the hareelv formation (katedralen member; fig. 3, table 1). the upper part of the sjællandselv-3 core overlaps with the lower portion of the sjællandselv-2 core; the uppermost level of the latter ties with the lower sjællandselv-1 core (see bjerager et al. 2018a, this volume). the succession comprises intercalated sandy deposits and mudrocks that range from laminated to more or less structureless. in general, the mudstone samples collected from the cores are rich in organic carbon, showing average values between 4.7 wt% and 6.0 wt% (table 3). the average petroleum potential varies from 7.9 kg/ton to 10.1 kg/ton, giving hi values in the range from 163 to 169 (fig. 6, table 3). on average, both tmax and the production index (pi) indicate pre-oil window maturity, but the average values mask widely varying values of both tmax and pi (fig. 6). only the sjællandselv-1 and jyllandselv-1 sample sets offer the possibility to establish the amount of non-reactive carbon and the average hydrogen index of the reactive fraction by cross-plotting toc versus s2 (dahl et al. 2004) using linear axes. datasets from other cores show too wide scatter to allow the construction of reliable trend-lines. using only samples showing s2 > 2 kg/ton and pi < 0.12 in order to avoid non-source samples, the jyllandselv-1 core yields an average toc of 7.7 wt%, of which 2.6 wt% will be inert and an average hydrogen index for the reactive kerogen of 314 (i.e. based on an average toc of 5.1 wt%). the sjællandselv-1 core yields an average toc, of 6.9 wt%, of which 3.3 wt% will be inert, and an average hydrogen index for the reactive kerogen of 241 (i.e. based on an average toc of 3.6 wt%; fig. 7). neither systematic palynofacies data nor organic petrographic data are available, but from the study of slides for palynological dating, it is known that the deposits are often very rich in woody terrestrial organic matter. bl ok el v1 bo re ho le 51 11 01 sjæ lla nd se lv1 bo re ho le 30 31 14 sjæ lla nd se lv2 bo re ho le 30 31 15 sjæ lla nd se lv3 bo re ho le 30 31 16 lo lla nd se lv1 bo re ho le 30 31 39 lo lla nd se lv2 bo re ho le 30 31 40 fa lst er el v1 bo re ho le 30 31 41 139 40 21 43 28 38 58 min 0.06 0.32 0.07 0.37 1.70 0.69 1.20 max 18.97 10.38 6.56 17.09 7.71 12.80 13.77 mean 5.94 5.58 4.71 6.02 3.89 3.98 4.78 min 305 391 364 363 420 415 415 max 473 447 441 446 434 432 436 mean 430 434 428 432 426 425 421 min 0.00 0.02 0.48 0.52 0.01 0.00 0.02 max 3.43 3.29 2.62 4.55 0.23 0.40 0.55 mean 1.31 1.32 1.43 1.86 0.06 0.07 0.14 min 0.00 0.00 0.34 0.66 0.74 0.03 0.63 max 42.91 30.29 14.14 21.72 7.81 5.43 20.51 mean 14.76 10.11 7.91 9.28 2.26 1.84 4.84 min 3 0 30 53 22 4 21 max 363 292 270 335 124 170 291 mean 218 169 163 166 59 48 92 min 0.00 0.05 0.06 0.08 0.01 0.00 0.01 max 0.70 0.56 0.64 0.80 0.05 0.18 0.18 mean 0.13 0.13 0.21 0.20 0.03 0.05 0.04 hi pi no. (samples) toc tmax s1 s2 jy lla nd se lv1 bo re ho le 30 31 42 83 1.63 11.30 7.56 426 451 439 0.28 3.13 1.68 1.82 27.15 15.55 69 262 196 0.05 0.15 0.10 jbk table 3 table 3. summary of petroleum potential data for borehole samples, all data sjællandselv-1: toc = 0.24 x s2 + 3.3 jyllandselv-1: toc = 0.32 x s2 + 2.6 gø02_13_139_jbk_blokelv 0 10 20 30 40 s2 (kg hc/ton rock) 0 4 8 12 16 to c (% ) jbk 7 fig. 7. linear plot of s2 vs. toc for samples from the sjællandselv-1 and the jyllandselv-1 cores; only samples showing s2 > 2 kg/ton and pi < 0.12 are used. the slope of the regression line represents the average hydrogen index of the reactive fraction of the kerogen. the intercept with the toc-axis represents the average proportion of non-reactive organic carbon present (see text for discussion). 95 blokelv-1 borehole the fully cored blokelv-1 borehole penetrated 233.8 metres of middle oxfordian to lower volgian deposits (fig. 3, table 1), consisting of c. 55% mudrocks, with the remainder being gravity-flow sandstones and sandstone injectites, plus four thin basaltic intrusions up to 2 m in thickness (bjerager et al. 2018b, this volume; larsen 2018, this volume). the blokelv-1 core covers the entire kimmeridgian succession, which seems to be without major hiati (alsen & piasecki 2018, this volume), and thus represents an important portion of the upper jurassic that is otherwise left more or less untested by outcrop fig. 8. crude oil bleeding from a belemnite cut by the blokelv-1 core (105.18 m depth). belemnite is c. 16 mm in diameter (parallel to lamination). 0 4 8 12 16 20 toc (wt%) 0 1 2 3 4 s1 (kg hc/ ton rock) 0 20 40 s2 (kg hc/ ton rock) 400 440 480 tmax (°c) 0 200 400 hydrogen index 0 4 8 12 16 20 tc (wt%) 220 200 180 160 140 120 100 80 60 40 20 0 m et re s b el ow re fe re nc e le ve l 0 2 4 6 8 ts (wt%) 0 0.4 0.8 1 production index jbk 9 fig. 9. blokelv-1 core, organic geochemical screening data vs. depth. from left to right: tc: total carbon. toc: total organic carbon. ts: total sulfur. s1, s2, hydrogen index, tmax, production index. shading indicates approximate intervals affected by magmatic intrusions (see text for discussion). 9696 jbk fig 10 gø02_13_128_jbk_blokelv immature oil condensate wet gas dry gas 0.1 1 10 100 toc (%) 0.1 1 10 100 excellent ex ce lle nt good g oo d fa ir po or poor 400 450 500 550 tmax (°c) 0 100 200 300 400 500 600 700 800 900 1000 h yd ro ge n in de x s2 (k g hy dr oc ar bo ns /to n ro ck ) pr od uc tio n in de x (s 1/ (s 1+ s2 )) 380 390 400 410 420 430 440 450 460 470 480 490 500 510 520 tmax (°c) 0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1 a b c staining or contamination low maturity conversion fig. 10. blokelv-1 core, organic geochemical screening data. a: tmax vs. hydrogen index. b: toc vs. s2. c: tmax vs. production index. 97 sampling and borehole drilling, including the prominent kimmeridgian sea-level highstand succession known from both the viking graben and east greenland (sneider et al. 1995; surlyk 2003). a total of 139 samples were analysed for petroleum potential (table 1). basaltic intrusions, at c. 102 m, 55 m, 28 m and 7 m (larsen 2018, this volume) resulted in the petroleum potential being more or less exhausted in a number of samples (table 3), and petroleum generation in close proximity to the sills was also observed. at 55.4 m, oil stains were present in a carbonate-lined fracture in the intrusion itself, and liquid oil was seen flowing from the phragmocone of a belemnite at 105.25m (fig. 8). in order to assess the overall petroleum potential of the succession, samples affected by heating from intrusions should be excluded, since intrusions are local phenomena; the following is thus based on a subset of the data (table 5) after all samples showing s2 < 2 and pi > 0.12 were excluded. the level of toc is highly variable, reaching values exceeding 18 wt%, with an average toc of nearly 7 wt% (table 3). an irregular, but generally increasing trend in toc extends from td to c. 170 m, followed by a gradual decreasing trend that extends to the uppermost parts of the borehole (fig. 9). the ts content shows a very irregular depth trend with initial values of 2–3 wt% from td to c. 120 m, at which point a steep increase begins, culminating at c. 8 wt% close to 70 m, from where an irregular decreasing trend is observed, falling to values close to 1 wt% near 10 m depth. the petroleum potential as represented by s2 is strongly variable, reaching values close to 43 kg/ton, with an average s2 of nearly 18 kg/ ton (table 5). ignoring the effects of the intrusions mentioned above, the average hydrogen index is fairly stable at close to 300 – slightly higher from c. 90 m to 10 m and slightly lower in the interval from td to c. 120 m. the effects of the intrusions are clearly seen in the hi, tmax and pi values (figs 9, 10c). excluding samples affected by heating by intrusions and non-source rock samples, the succession is seen to contain primarily oil-prone type ii or mixed oiland gas-prone type ii/iii kerogen, with an average hi of 328 for the active kerogen and an average inert organic carbon content of only 0.95% (fig. 11). the thermal maturity of the succession is close to the oil window and the increasing trend in pi points to incipient generation, despite pi in absolute terms remains low (<0.12). the thermal maturity indicators consistently show an increasing trend with depth (fig. 12a, table 6), with the start of the oil-generative window being found somewhere between 100 m and 160 m, indicating that considerable uplift has occurred. the magmatic intrusions have had a major impact on the thermal maturity of the sedimentary rocks immediately adjacent to them, but the effects are local, restricted to an aureole around the intrusions amounting to only three to five times their thickness. an example of this is shown in fig. 13, where the changes in hydrogen index (hi) and production index (pi) as well as in organic petrography are illustrated. both above and below the intrusion, the hi decreases steadily approaching the intrusion as the petroleum potential is realised. the pi shows a parallel increase due to petroleum generation, but close to the intrusion, the pi shows a rapid decrease due to thermal destruction of the petroleum and generation of gaseous components that tend to escape. under the microscope, the changes are manifest in gradually decreasing fluorescence and, closest to the intrusion, generation of pyrobitumen, coke and natural chars. a number of different parameters, in addition to the screening data, show a gradual change with depth, with a marked shift in the trends at c. 90 m (fig. 12b, table 6). the pristane/phytane ratio shows a steady decrease from c. 6 at td to values between 1 and 3 in the upper 90 m of the succession. the isohopane ratio (nytoft 2011) decreases from more than 0.12 at td to an average of 0 10 20 30 40 50 4 0 8 12 16 20 toc = 0.328 x s2 + 0.95 to c (% ) jbk 11 s2 (kg hc/ton rock) fig. 11. blokelv-1 core. linear plot of s2 vs. toc; selected data only, s2 > 2 kg/ton and pi < 0.12. the slope of the regression line represents the hydrogen index of the reactive fraction of the kerogen. the intercept with the toc-axis represents the average proportion of non-reactive organic carbon present (see text for discussion). 0 0.04 0.08 0.12 0.16 isohopane ratio 0 0.2 0.4 0.6 0.8 1 bisnorhopane ratio 2 4 6 8 10 total s30 (%) -34 -32 -30 -28 -26 δ13c (saturates) 1 2 3 4 5 6 pristane/phytane 220 200 180 160 140 120 100 80 60 40 20 0 0 0.05 0.1 0.15 homohopane ratio 0 0.2 0.4 0.6 0.8 production index 0.3 0.4 0.5 0.6 s29 20s/(20s+20r) 0.3 0.4 0.5 0.6 s29 αββ/(αββ+ααα) 0.6 0.7 0.8 0.9 1 total dia/(dia+regular) 400 420 440 460 480 tmax (°c) 220 200 180 160 140 120 100 80 60 40 20 0 m et re s b el ow re fe re nc e le ve l m et re s b el ow re fe re nc e le ve l 0.4 0.6 0.8 1 ro (%)a b jbk 12 fig. 12. a: blokelv-1 core, thermal maturity indicators vs. depth. from left to right: tmax (from rock-eval type screening). production index (from rock-eval type screening). ro (%): vitrinite reflectance. s29 20s/(20s + 20r): regular c29 sterane ααα20s/(ααα20s+ααα20r) isomer ratio. s29 αββ/(αββ+ααα): regular c29 sterane αββ/(αββ+ααα) isomer ratio. total dia/(dia + regular): sum c27–29 diasteranes/ (sum c27–29 diasteranes + sum c27–29 regular steranes). red symbols indicate oil sample (stain) data. shading indicates approximate intervals affected by magmatic intrusions (see text for discussion). b: blokelv-1 core, facies indicators versus depth. from left to right: pristane/phytane ratio. isohopane ratio (nytoft 2011). homohopane ratio: h35/(sum h31–35). bisnorhopane ratio: h28/h30. total s30 (%): percentage of c30 steranes relative to total c27–30 steranes. δ13c (saturates): stable carbon isotope δ13c of saturated extract fractions. red symbols indicate oil sample (stain) data. shading indicates approximate intervals affected by magmatic intrusions (see text for discussion). 99 c. 0.08 in the upper 90 m of the succession, whereas the homohopane ratio increases from a value close to 0.02 to an average of c. 0.08 over the same interval. the bisnorhopane ratio shows strong variation but with a rapid increase from a near-zero value at td to c. 1 over the lowermost 60–70 m. the percentage of marine c30 steranes (n-propylcholestanes) increases from values close to 3% near td to values close to 8% in the upper c. 90 m of the succession, whereas δ13c shows a trend from relatively less depleted isotopic compositions (c. –27‰) to more strongly depleted compositions (c. –30‰) over the same interval. this feature is even more clearly expressed when the isotopic composition of the total extract is plotted against the pristane to phytane ratio (fig. 14), where the broad right-to-left trend indicated represents the overall up-section development. where applicable, data on oil stains collected at 55.4 m and 105.18 m are shown together with rock data. the deeper sample (105.18 m) shows a near-identical composition relative to nearby rock samples in all parameters considered, whereas the shallower sample (55.4 m) shows slight to moderate deviation in some parameters (fig. 12). characteristic fingerprints of samples representing the upper (sample 16827, 80.77 m) and lower (sample 16850, 205.84 m) parts of the drilled succession are shown in fig. 15, and for comparison corresponding data on the two oil stain samples are shown in fig. 17. a common feature of all samples is low to moderate proportions of rearranged hopanes such as neohopanes and diahopanes, whereas the proportions of rearranged steranes are generally somewhat higher. relative to samples of the upper part of the succession, samples of the lower part show much higher proportions of longchain (‘waxy’) n-alkanes and a much higher pristane/ phytane ratio. the entire succession is remarkably low in tricyclic terpanes, but the samples from the deeper part of the succession clearly contain even less than samples 0 100 200 300 40050 150 250 350 hydrogen index (mg hc/g toc) sill 34 32 30 28 26 2 3 1 24 22 d ep th (m ) 0 0.2 0.4 0.6 0.80.1 0.3 0.5 0.7 production index (s1/(s1+s2) hydrogen indexproduction index 1 2 3 3 3 natural chars coke pyrobitumen gø02_13_138_jbk_blokelv jbk fig 13 decreasing fluorescence towards sill (samples 1–3) due to enhanced maturity fig. 13. an example of the effects of magmatic intrusions on the thermal maturity and petroleum potential of adjacent deposits. the petroleum potential of deposits adjacent to magmatic intrusions was rapidly realised, as shown by decreasing hi. the pi increases sharply due to petroleum generation, but closest to the intrusion, the pi decreases again due to cracking and thermal destruction of petroleum to form gaseous components that tended to escape. the fluorescence-intensity of the liptinite group decreased as petroleum was generated (samples 1–3, see cross-plot) and high temperatures immediately adjacent to the intrusions led to the formation of pyrobitumen and natural chars (see sample 3, photos right), whereas some sedimentary coal-particles were transformed to coke. 100100 from the upper part. samples from the upper part of the succession often contain high proportions of 28,30-bisnorhopane, and extended hopanes, whereas samples from the deeper part only contain minor proportions of 28,30-bisnorhopane as well as of extended hopanes. c30 n-propylcholestanes are scarce in the lower portion of the succession but increase rapidly in abundance upwards as the overall distribution of steranes changes from a strong predominance of c29-steranes to near-equal proportions of c27, c28 and c29-steranes (fig. 15, table 6). the variation of kerogen composition with depth demonstrated by the various organic geochemical parameters mentioned above are clear in the petrographic composition of the kerogens present (fig. 16). evidently, the kerogen of the deeper part of the succession is predominantly terrestrial, being composed of vitrinite and inertinite, with only minor proportions of detrital liptinite, presumably alginite, and fluorescent amorphous organic matter. conversely, the kerogen of the upper part of the succession is predominantly marine, being composed of alginite and fluorescent amorphous organic matter, with only minor proportions of terrestrial constituents such as vitrinite and inertinite. the shallower oil stain sample collected at 55.4 m, close to the most organic-rich part of the source-rock succession, shows clear signs of biodegradation leading to a marked reduction in the abundance of normal alkanes and a corresponding enhancement of more complex moieties, including hopane and sterane biological markers, which may even be discerned in the gc-fid data (fig. 17). in addition, a prominent envelope of unresolved complex components is observed. although individual biological marker ratios may deviate slightly from expected values, the overall characteristics of the sample correspond well to those of nearby source-rock samples. the deeper oil stain sample collected at 105.18 m, c. 25 m below the initiation of the change in geochemical characteristics with depth, is largely unaffected by biodegradation but may have suffered minor evaporative losses in the lower boiling point range. in all measured parameters, this sample shows close correlation to neighbouring source-rock samples. within the entire succession, it occupies an intermediate position, showing characteristics between those of the shallower and deeper parts of the source rock succession (figs 12, 14, 15, 17, table 4). discussion although stratigraphic datings are not available from all outcrop samples, existing evidence indicates that these samples and the ‘old boreholes’ represent the upper part of the olympen formation and the lowermost and the topmost parts of the katedralen member (hareelv formation), whereas the central parts of the hareelv formation are left essentially unsampled (fig. 3). the blokelv-1 core provides full coverage of the stratigraphic intervals that were not represented in other sample material as well as of duplicating of the succession that was represented by older core and outcrop samples; note, however, that the blokelv-1 borehole did not penetrate the base of the hareelv formation. in the northern north sea region, deposition of organic-rich shales continued into the early cretaceous (ryazanian), with the richest and most oil-prone deposits being present in the upper –32 –31 –30 –29 –28 –27 –26 –25 –24 –23 δ13c (total oil/extract) 0 1 2 3 4 5 6 pr ist an e/ ph yt an e ra tio marine shale and palaeozoic carbonate deltaic mesozoic carbonate up section gø02_13_131_jbk_blokelv jbk fig 14 sts fig. 14 blokelv-1 core, stable carbon isotope δ13c of total extract fractions (calculated by combining extract composition and isotope data for individual fractions) vs. pristane/phytane ratio. red symbols indicate oil sample (stain) data. plot after chung et al. (1992). from the base of the core upwards, there is a general right-to-left trend from high pristane/phytane and undepleted δ13c to lower pristane/phytane and increasingly depleted δ13c, mirroring a facies change from terrestrially influenced kerogen to increasingly marinedominated kerogen (see text for discussion). facing page: fig. 15. biological marker data. characteristic fingerprints of samples representing the upper (sample 16827, 80.77 m) and lower (sample 16850, 205.84 m) parts of the drilled succession. gc-msms parent–daughter traces for steranes and hopanes represent the sum of five and nine parent–daughter transitions, respectively. 101 time sig na l time sig na l time sig na l time sig na l time sig na l time sig na l time sig na l time sig na l pri sta ne pri sta ne ph yt an e ph yt an e gc-ms(sim) m/z 191 gc-ms(sim) m/z 191 gc-fid gc-fid gc-ms-ms parent ➝ daughter hopanes gc-ms-ms parent ➝ daughter hopanes gc-ms-ms parent ➝ daughter steranes gc-ms-ms parent ➝ daughter steranes ts tm h28 h29 h30 h31 (s +r ) h32 (s +r ) h33 (s +r ) h34 (s +r ) h35 (s +r ) d 30 m 29 m 30 ts tm h28 h29 h30 h31 (s +r ) h32 (s +r ) h33 (s +r ) h34 (s +r ) h35 (s +r ) d 30 m 29 m 30 s29 αα αs αα αs αα αs αα αr αα αr αα αr αα αr αβ βr αβ βr αβ β αβ β αβ βs αβ βs s30 s29 s30 d26d26 d27 d27 d26 d26 d27 d27 sample 16827, 80.77 m sample 16850, 205.84 m t 23 t 24 t 25 t 26 te 24 t 25 t 26t 24 te 24 nc 20 nc 25 nc 30 nc 35 nc 40 nc 15 nc 20 nc 25 nc 30 nc 35 nc 40 nc 15 jbk 15 102102 volgian – ryazanian succession (e.g. ineson et al. 2003 and references therein; justwan et al. 2006a). in contrast, most of the corresponding volgian–ryazanian succession in jameson land is sand-rich, although a few mudstone samples with rather modest petroleum potential have been collected from the ryazanian hesteelv formation in southern jameson land. the hesteelv formation, which was deposited in an incised valley during the initial phase of regional flooding in the ryazanian (surlyk 1973; surlyk et al. 1973; f. surlyk personal communicasample depth (m) pr/ph bnh/ h30 29ts d30 ihr hhi s27 (%) s28 (%) s29 (%) s30 (%) s29 (s/s+r) s29 ββ/ (ββ+αα dia/ (dia+ reg) δ13c (sat) 511101-107 12.62 2.55 0.17 0.20 0.07 0.06 0.09 28.2 30.1 34.3 7.5 0.35 0.33 0.72 –30.5 511101-110 16.42 2.36 0.92 0.17 0.08 0.06 0.12 29.3 28.8 33.9 8.1 0.37 0.33 0.67 511101-113 20.15 2.25 0.11 0.19 0.07 0.07 0.09 28.6 30.7 32.6 8.0 0.36 0.34 0.71 –30.6 511101-8 23.59 2.01 0.32 0.13 0.05 0.07 0.07 30.4 32.9 28.4 8.3 0.41 0.33 0.66 –30.2 511101-117 24.26 1.95 0.24 0.12 0.05 0.07 0.07 30.6 32.1 28.5 8.8 0.43 0.34 0.66 511101-139 24.68 1.88 0.11 0.13 0.05 0.07 0.06 31.5 31.3 28.6 8.6 0.46 0.36 0.66 511101-118 25.15 1.91 0.08 0.10 0.06 0.07 0.07 32.5 31.3 27.3 8.9 0.50 0.48 0.76 –29.4 511101-137 25.24 1.92 0.01 0.11 0.06 0.07 0.06 34.1 31.6 26.1 8.2 0.54 0.58 0.83 –29.1 511101-119 25.76 1.77 511101-135 25.88 1.93 0.24 0.13 0.07 0.09 0.07 34.8 31.0 26.6 7.6 0.43 0.44 0.82 511101-121 26.23 1.84 0.40 0.11 0.08 0.07 0.09 33.1 30.3 28.4 8.1 0.40 0.35 0.67 –29.0 511101-122 27.11 1.86 0.03 0.13 0.10 0.06 0.10 33.9 30.0 28.4 7.7 0.40 0.36 0.69 511101-141 27.46 1.77 0.40 0.11 0.08 0.07 0.08 33.7 30.4 28.4 7.5 0.42 0.39 0.73 –32.3 511101-123 28.33 1.62 511101-144 28.69 2.64 0.03 0.08 0.05 0.09 0.04 31.1 28.9 33.8 6.2 0.55 0.58 0.83 –28.9 511101-124 29.24 1.84 0.20 0.11 0.05 0.07 0.08 29.8 32.3 29.2 8.7 0.49 0.43 0.71 511101-125 32.10 1.92 0.49 0.13 0.06 0.06 0.09 30.8 32.9 28.7 7.5 0.42 0.34 0.65 –30.1 511101-12 32.78 1.96 0.50 0.11 0.05 0.06 0.10 30.0 32.1 29.7 8.3 0.41 0.33 0.68 –30.4 511101-146 50.27 2.06 0.44 0.14 0.07 0.08 0.09 30.4 28.3 32.9 8.4 0.45 0.34 0.66 511101-150 64.91 2.10 0.66 0.15 0.07 0.06 0.09 30.9 29.9 31.2 8.0 0.46 0.36 0.68 –30.0 511101-24 68.77 1.97 0.38 0.14 0.07 0.07 0.09 33.1 27.9 31.8 7.2 0.47 0.36 0.66 511101-28 80.77 2.14 0.56 0.13 0.06 0.07 0.09 31.9 28.1 31.1 8.9 0.48 0.37 0.72 –29.7 511101-164 82.14 1.89 0.18 0.16 0.07 0.07 0.10 34.3 27.1 29.8 8.7 0.49 0.40 0.72 511101-170 89.22 2.42 0.10 0.11 0.06 0.07 0.08 29.2 25.8 38.8 6.2 0.49 0.40 0.74 –29.3 511101-32 92.74 2.36 0.11 0.12 0.06 0.08 0.08 29.5 25.6 39.1 5.8 0.50 0.40 0.74 511101-174 93.02 2.16 0.27 0.11 0.06 0.08 0.08 30.3 25.3 39.0 5.4 0.50 0.41 0.74 –29.1 511101-175 97.42 1.99 0.07 0.05 0.05 0.08 0.05 35.4 27.1 32.8 4.7 0.55 0.57 0.80 –28.5 511101-176 98.48 2.67 0.40 0.08 0.04 0.10 0.05 32.4 24.4 38.2 5.0 0.46 0.44 0.71 511101-177 103.98 2.74 511101-36 104.79 2.43 0.09 0.05 0.03 0.09 0.05 33.3 25.9 36.1 4.6 0.53 0.55 0.77 –28.1 511101-178 105.72 2.50 0.21 0.09 0.03 0.08 0.07 28.6 24.8 41.2 5.4 0.51 0.43 0.72 –28.6 511101-179 106.59 2.71 0.38 0.10 0.05 0.01 0.07 29.7 26.1 38.4 5.8 0.50 0.44 0.74 –28.7 511101-186 117.73 2.60 0.77 0.13 0.07 0.08 0.09 37.7 23.9 31.3 7.1 0.50 0.42 0.75 511101-195 142.84 3.08 0.40 0.09 0.08 0.09 0.06 28.1 21.7 43.6 6.5 0.52 0.48 0.75 511101-203 160.80 3.52 0.95 0.08 0.06 0.09 0.05 37.0 17.8 40.9 4.3 0.50 0.52 0.76 –27.1 511101-60 176.77 3.83 0.43 0.13 0.07 0.08 0.05 29.6 20.7 44.5 5.2 0.49 0.53 0.82 511101-220 205.84 3.66 0.09 0.11 0.07 0.11 0.04 20.8 18.9 56.2 4.1 0.53 0.54 0.78 –27.3 511101-228 232.05 5.77 0.01 0.09 0.09 0.12 0.02 18.2 19.3 59.9 2.7 0.50 0.55 0.84 oil stain 105.18 2.31 0.12 0.04 0.03 0.08 0.08 31.0 25.4 38.3 5.2 0.53 0.47 0.75 –28.5 oil stain 55.4 2.05 0.63 0.09 0.05 0.06 0.14 31.2 25.8 37.1 5.9 0.47 0.37 0.72 –28.5 table 4. gas chromatography, biomarker and isotopic data on core extracts and oil stains pr/pr: pristane/phytane ratio. bnh/h30: 28,30-bisnorhopane/hopane ratio. ihr: isohopane ratio (nytoft 2011). hhi: homohopane index (h35/σ(h31-h35)). s27 (%), s28 (%), s29 (%), s30 (%): distribution of total c27–c30 steranes (diasteranes + regular steranes), normalised to 100%. s29 (s/(s+r)): c29 sterane 20s/(20s+20r) isomer ratio. s29 ββ/(ββ+αα): c29 sterane αββ/(aββ+ααα) isomer ratio. dia / (dia + reg): total c27–30 diasteranes /(total c27–30 diasteranes + total c27–30 regular steranes). δ13c (sat): stable carbon isotopic ratio, saturated fraction. jbk table 4 103 tion, 2014), represents the youngest mesozoic deposits in jameson land. despite the muddy nature of the lower crinoid bjerg member of the hesteelv formation, the highly oil-prone shales known from the uppermost volgian – ryazanian succession in other parts of the greater north atlantic region seem not to have been deposited in jameson land although they are recorded farther north in north-east greenland. outcrop samples outcrop samples generally show pre-oil window thermal maturity, except for samples from the fortet and katedralen outcrops affected by widespread basaltic intrusions, and are rich in organic matter although their levels of toc show wide variation. in absolute terms, as represented by the s2-parameter, their petroleum potential is generally good or even excellent, but when normalised by toc to calculate the hydrogen index, the kerogen is classified as primarily gas-prone type iii or gas/oil-prone type ii/iii. hence, compared to correlative stratigraphic intervals in the greater north atlantic area, the apparent petroleum potential is surprisingly poor. this may in part be attributed to the high proportions of oxidised and terrestrial organic matter observed in palynological preparations, which may serve to ‘downgrade’ the hydrogen index by adding large amounts of organic carbon that do not participate in petroleum generation. an additional cause may be the very large concentrations of finely disseminated pyritic sulfur present in the shales. very few analyses of ts have been made on outcrop samples, but systematic analyses of the blokelv-1 core (fig. 9) indicate that fresh/unweathered core samples represent the original values of the outcrop samples. weathering of pyrite will ultimately generate sulfuric acid, which, by its oxidising nature, attacks kerogen and reduces its petroleum potential. extensive weathering of pyrite is evident on outcrop faces that are commonly encrusted with greenish-yellow jarosite, a potassium-iron sulphate-mineral (kfe3+3(oh)6(so4)2) characteristically formed by pyrite weathering (fig. 18). hence, data on the petroleum potential derived from surface/outcrop samples are not representative if the rocks contain large amounts of finely disseminated pyrite. ‘old boreholes’ these core samples show strongly variable petroleum potential, but only represent a restricted stratigraphic interval since the lower and middle kimmeridgian intervals were not penetrated. the majority of the samples can be classified as either gas-prone, containing predominantly type iii kerogen, or as non-source deposits containing predominantly inert kerogen. however, a large proportion of the samples shows hi values of 200–300 or higher and may be classified as gas/oil-prone to oil prone. on average, both tmax and the production index (pi) indicate pre-oil window maturity, but the average values mask widely varying values of both tmax and pi, suggesting localised maturation and petroleum generation caused by heating from the abundant intrusions present in the area. although it is impossible to make a straight comparison due to the stratigraphical and lithological differences, it seems that core samples do show higher average petroleum potential than outcrop samples due to weathering of the latter. in the sjællandselv-1 and the jyllandselv-1 cores, assessment of the quality of the reactive kerogen fraction can be ascertained by the linear trendline shown by the toc vs. s2 plot (fig. 8). this shows that although there is a very large fraction of inert carbon that does not take part in petroleum generation, the reactive fraction of the kerogen is generally oil and gas to oil prone. this is in close accordance with results from the blokelv-1 core, although the average proportion of inert carbon is significantly lower. blokelv-1 petroleum potential and organic facies variations due to its continuous stratigraphy from the middle oxfordian to the lower volgian, the blokelv-1 cored section serves as a reference for all other sampling locations, both outcrops and ‘old boreholes’. moreover, it includes the lower to middle kimmeridgian interval which, apart from being largely untested by other sampling, constitutes a more prolific petroleum source succession than any other interval known in the upper jurassic of jameson land. based on petroleum geochemical data, it is clear that the succession penetrated by the blokelv-1 borehole is divided into two subunits (fig. 9). a lower unit, extending from td to c. 110 m, is characterised by generally high levels of toc, moderate sulfur content, and high s2-yields, resulting in hi-values slightly below 300 on average, showing a clear upward-increasing trend. these 104104 a b c d e f g h 105 deposits accumulated during overall rising sea level (bjerager et al. 2018b, c, this volume). the transition to the overlying unit is obscured by an intrusion which has affected the petroleum potential of the adjacent succession both above and below. by comparison, the upper unit from c. 90 m to c. 10 m is characterised by fairly high levels of toc, predominantly oil-prone kerogen and very high levels of sulfur, and slightly higher average hi. the unit shows a fairly clear maximum in all parameters over the interval from c. 80 m to c. 60 m. this interval was deposited during a prominent sea-level highstand (bjerager et al. 2018b, c, this volume). the drilled succession thus records changes in depositional conditions in response to a gradually rising sea level from td to c. 90 m, culminating in the interval 90–60 m; the analytical data suggest a weak regressive trend in the uppermost 60 m of the cored succession. although the general depositional environment is marine and strongly oxygen-restricted throughout the succession, the kerogen shows a temporal change from marine with a prominent terrestrial component to compositions with successively less terrestrial organic matter, culminating in the interval 90–60 m (fig. 16). this trend is also evident from biomarker and isotopic data (figs 12b, 14), notably pristane/phytane ratios, the proportion of marine c30 n-propylcholestanes and δ13c, but also isohopane and homohopane ratios, and even raw data (fig. 15). hence, a clearly decreasing trend in pristane/phytane ratio is observed upwards through the succession (fig. 12b). a similar, but a less pronounced trend is observed in the isohopane ratio, which is a very sensitive indicator of terrestrial organic contributions (nytoft 2011). upward-increasing trends are observed in the homohopane ratio, pointing to increasingly oxygendeficient depositional environments, and in the percentage of c30 n-propylcholestanes, pointing to increasing proportions of marine organic matter in the kerogen. stable carbon isotopic ratios show a clear trend from less depleted to more depleted isotopic compositions, suggesting changing proportions of terrestrial relative to marine kerogen in favour of the latter. conversely, the 28,30-bisnorhopane ratio does not show any clear trends, but the abundance of 28,30-bisnorhopane is remarkable when compared to equivalent deposits in the north sea region, where high levels of 28,30-bisnorhopane are characteristic of the uppermost, highly prolific volgian– ryazanian part of the succession whereas its abundance in the deeper parts, for instance in the kimmeridgian succession, is often low (e.g. ineson et al. 2003). oilstains recorded at 55.4 m and 105.18 m show biomarker fingerprints reflecting their stratigraphic position in the succession, in that they show characteristics very similar to the shales adjacent to the intrusion in or near which they occur (figs 12b, 14, 17). the deeper stain shows a near-perfect match to nearby samples whereas the shallower stain shows minor deviation in some parameters, probably related to its biodegraded nature (fig. 17). despite the changes in organic facies documented by a number of independent parameters, oil-prone petroleum source rocks are present throughout the entire drilled succession. moreover, as indicated by the hydrogen index, the difference in petroleum potential between the more terrestrial lower part and the upper more marine part is trifling. this unexpected observation may have to do with the balance between inert and labile sedimentary organic matter, where the latter can be expected to be ‘diluted’ by the former, which will then serve to drag down the hydrogen index. conceivably, the autochthonous marine organic matter is diluted by allochthonous inert material. this effect will be more pronounced the lower the toc, if the autochthonous contribution is more or less constant. however, based on the geochemical signature of the oil-stains found, it is clear that even the terrestrial fraction of the kerogen in the lower part of the succession is oil prone since both stains show clear relationships to the shales immediately adjacent to them. facing page: fig. 16. examples of kerogen composition in the upper (a–d) and lower (e–h) parts of the drilled succession. polished block preparations, reflected light microscopy, horizontal edge of photographs is c. 300 micron. photos a–d from sample 16827 (81 m). a: finely laminated mudstone, abundant framboidal pyrite (bright white) and vitrinite (grey). b: same field as a, fluorescence-inducing blue light. bright yellow alginate, brownish-yellow fluorescing amorphous organic matter. c: finely laminated mudstone, abundant framboidal pyrite (bright white), small vitrinite particles (grey). d: same field as c, fluorescence-inducing blue light. bright yellow alginate, brownish-yellow fluorescing amorphous organic matter. photos e–h from sample 16838 (213 m). e: mudstone, containing abundant vitrinite (grey) and inertinite (white), minor amounts of framboidal pyrite (bright white). f: same field as e, mudstone, scattered brownish-yellow fluorescing amorphous organic matter and minute detrital liptinite, presumably alginite. g: mudstone, containing abundant vitrinite (grey) and inertinite (white), minor amounts of framboidal pyrite (bright white). h: same field as g, mudstone, scattered brownish-yellow fluorescing amorphous organic matter and minute detrital liptinite, presumably alginite. 106106 time si gn al time si gn al time si gn al time time time time time si gn al si gn al si gn al si gn al si gn al pr ist an e pri sta ne ph yta ne ph yt an e nc 20 nc 25 nc 30 nc 35 t 25 t 26 t 25 t 24 te 24 t 25 t 26t 23 t 24 te 24 gc-ms(sim) m/z 191 gc-ms(sim) m/z 191 gc-fid gc-fid ts tm h28 h29 h30 h31 (s +r ) h32 (s +r ) h33 (s +r ) h34 (s +r ) h35 (s +r ) d 30 m 29 m 30 ts tm h28 h29 h30 h31 (s +r ) h32 (s +r ) h33 (s +r ) h34 (s +r ) h35 (s +r ) d 30 m 29 m 30 s29 s30 αα αs αα αs αα αs αα αr αα αs αα αr αα αr αα αr αβ β αβ β αβ βr αβ βs αβ βr s29 s30 d26d26 d27 d27 d26 d26 d27 d27 sample 16053, 55.4 m sample 16764, 105.18 m gc-ms-ms parent ➝ daughter hopanes gc-ms-ms parent ➝ daughter hopanes gc-ms-ms parent ➝ daughter steranes gc-ms-ms parent ➝ daughter steranes nc 15 nc 15 nc 20 jbk 17 107 hence, the above issue is not solely a matter of the balance between labile marine kerogen and more or less inert autochthonous kerogen, but includes qualitative changes in the labile kerogen as well. systematic organic petrographic data may perhaps shed some light on this matter, but in the absence of such data no acceptable explanation can be offered. blokelv-1 borehole thermal maturity the magmatic intrusions present within the drilled succession have clearly had a significant impact on the thermal maturity of the shales immediately adjacent to them (fig. 13), where the petroleum potential is rapidly realised, as shown by decreasing hi. approaching the intrusion, the pi increases sharply recording petroleum generation, but immediately adjacent to the intrusion, the pi decreases in response to cracking and thermal destruction of petroleum to form gaseous components that tend to escape. the fluorescence intensity of liptinite-group macerals decreases as petroleum is generated and high temperatures immediately next to the intrusions led to the formation of pyrobitumen and natural chars, whereas sedimentary coal-particles may form coke. these effects are clearly local, however, and have little influence on the overall level of thermal maturity of the section. in general, the level of thermal maturity of the blokelv-1 succession is within the early stages of the oilgenerative window, with incipient petroleum generation actually taking place, and in the vicinity of magmatic intrusions the petroleum potential has been more or less realised as evidenced by the presence of oil-stains. independent thermal maturity indicators such as tmax from rock-eval type pyrolysis, vitrinite reflectance (%ro) and biomarker data such as sterane isomerisation ratios and the diasterane ratio are all in perfect agreement, although the latter is also strongly facies-dependant (fig. 12a, tables 4, 5, 6). a remarkable feature is the clearly increasing trend with depth observed in all the parameters mentioned, with the transition to the ‘oil-window’ as defined by tmax > 435°c and ro >0.55% located somewhere in the interval between 120 m and 160 m. a clear gradient is unexpected over such a narrow depth interval (c. 234 m). it is evident that a significant thickness of overburden has been removed by erosion since the deposits were at their maximum depth of burial. it is estimated that 2–3 km of cretaceous-age sediments and palaeogene-age volcanics have been removed due to neogene uplift and erosion (mathiesen et al. 1995, 2000; green & japsen 2018, this volume). hence, with a removed section amounting to 2–3 km, the upper jurassic succession was close to or within the oil-generative window at its maximum depth of burial, given a modest enhancement of the geothermal gradient as is to be expected from the volcanic activity in the area at the time of maximum burial. moreover, the abundance of sands intercalated with the shales can conceivably be expected to have served to assist thermal maturation by facilitating fluid movement and thus overcoming time-lag effects caused by differences in thermal conductivity. the processes associated with thermal maturation and the conversion of kerogen into petroleum are not linear but include several ‘thresholds’, which in coalpetrographic nomenclature are known as ‘coalification jumps’ (see discussion in taylor et al. 1998), and defined by initiation or rapid changes in the rates of certain processes. incipient petroleum generation or the start of the ‘oil window’ coincides with the first ‘coalification jump’ where a number of processes associated with petroleum generation accelerate, causing the non-linearity of matufacing page: fig. 17. biological marker data. characteristic fingerprints of oilstain samples, sample 16053 (55.4 m) and sample 16764 (105.18 m). gc-ms-ms parent–daughter traces for steranes and hopanes represent the sum of five and nine parent–daughter transitions, respectively. fig. 18. outcrop of the katedralen member shales near the blokelv-1 drill site, showing widespread coating by jarosite (orange–yellow) formed from weathering of pyrite. 108108 ration profiles observed in exploration wells entering or penetrating the oil generative window. the observed thermal maturity of the blokelv-1 succession straddles the transition to the oil-generative window, but generation ceased due to uplift and erosion, removing the overlying succession, hence the unusual gradient in particular at depths greater than c. 120 m. blokelv-1 core: calculation of ultimate expulsion potential (kinex™ modelling) ultimate expulsion potential modelling based on the succession penetrated by the blokelv-1 borehole calls for caution since both in terms of organic facies variations and thickness of the succession modelled, the basin is not very well constrained. the blokelv-1 borehole is merely a point-source of information, and caution should be exer-50 14000 28 30 32 34 36 38 40 42 44 12000 10000 8000 6000 4000 2000 0 0 1 2 vitrinite reflectance – llnl (%ro) c um ul at ive g lr e xp el le d (s cf /st b) 3 4 5 50 100 150 200 250 300 0 0.5 1 uep (mmboe/km2) th ick ne ss (m ) 0 0 0 0.2 0.4 0.6 0.8 1 1 2 vitrinite reflectance – llnl (%ro) 3 4 5 0 1 2 vitrinite reflectance – llnl (%ro) 3 4 5 cumulative hc6+ expelled (/mass) cumulative hc1–5 expelled (/mass) c um ul at ive h c 6+ e xp el le d (/m as s) c um ul at ive h c 6+ a pi e xp el le d (/m as s) a b c d hc6+ hc1–5 jbk 19 cumulative glr expelled (scf/stb) total uep = 61.3981 mmboe/km2 thickness = 233.94 m cumulative hc6+ gravity expelled (api) fig. 19. calculation of the ‘ultimate expulsion potential’ (uep) by kinex™-modelling results. a: ultimate expulsion potential (uep) vs. depth showing an oiland gas-prone mudstone section with a total uep of about 61.4 mmboe/km2, but with a potential for generation of approximately twice as much gas as oil. b: cumulative expelled c6+ hydrocarbons (oil) and c1–5 hydrocarbons (gas). oil generation occurs from about 0.7–1.1% vitrinite reflectance. c: cumulative expelled c6+ oil gravities (api) if the mudstones passed through the oil window. the api range suggests generation of medium to light oil. d: cumulative expelled glr (gas-liquid ratio). 109 cised in extrapolating the modelling results to a more regional scale, although for simple comparison with similar successions in the north atlantic region, such modelling may still be meaningful. the hi and toc values show that the petroleum generation potential varies through the blokelv-1 well, which reflects organofacies variations in the marine mudstones. a method to monitor these variations in a more quantitative way is to estimate the potential charged fluid phases (oil and gas) through the mudstone section by calculating the ultimate expulsion potential (uep; mmboe/km2). uep displays the ultimate expellable volume of hydrocarbons in million barrels of oil equivalents (mmboe) per km2. the kinex™ tool (zetaware™) was used to calculate the uep and to carry out simple maturation modelling of the marine mudstones using a default heating rate of 2°c/ma and the kinetics of organofacies b (marine clay-rich shale) as defined by pepper & corvi (1995). a depth profile of the uep calculated from the measured toc and hi values is shown in fig. 19a. calculation of uep is affected by the thickness of the individual sample intervals. this means that a sample representing a comparatively thick interval will yield a disproportionate large uep compared to regular thin interval samples. a profile of uep may therefore be influenced by irregular core sampling intervals resulting in contrasting thicknesses used for the calculations. to overcome this problem and obtain a more representative uep profile through the mudstone section in blokelv-1, the c. 234 m thick section was subdivided into intervals each c. 1–2m thick. most sample intervals were within this range, thus only a limited number of samples were affected. for example, a 2.4 m thick interval was divided into two subsections of 1.2 m, each interval being assigned the toc and hi values of the corresponding original sample. the blokelv-1 well drilled a number of basaltic intrusions and these intervals were assigned toc and hi values of zero. the total (cumulative) uep is c. 61.4 mmboe/km2, but varies through the succession in response to the varying toc and hi values. this number is in good agreement with estimates of the potential of other kimmeridge clay formation (kcf) equivalent successions in the north sea and siberia (demaison & huizinga 1991) even though the blokelv-1 core represents only a fraction of the total thickness of the kcf-equivalent succession. the igneous intrusions and the adjacent mudstones with overcooked kerogen are clearly identified by the gaps in the profile with no potential. sandstone beds have not been subtracted from the uep calculation, and the total uep is thus a maximum estimate, since the kinex™ tool cannot accommodate the full complexity of the interbedded mudstone–sandstone succession. the lower part of the cored section has a fluctuating oil potential, but a higher capacity to generate and expel gas, so although the mudstones have a propensity to generate both oil and gas, they are primarily gas-prone. the oil generation potential increases up-section and is at a maximum in the interval 175–130 m, a trend followed by the gas generation potential (fig. 19a). above 130 m, the oil generation potential decreases, attaining a relatively stable level from 130 m to 70 m, above which point it decreases progressively up-section. the uop (ultimate oil potential) is c. 20 mmboe/km2 whereas the ugp (ultimate gas potential) is c. 41.4 mmboe/km2, i.e. a substantially higher proportion of gas than oil will be generated. the kinex™ model predicts that on passing through the oil window, the mudstones will generate c6+ hydrocarbons from c. 0.7%ro to slightly above 1%ro with a cumulative api ranging from c. 29° to c. 44° (figs 19b, c). the charge up to a vitrinite reflectance of about 1.1–1.2%ro will be ‘black’ to ‘volatile’ oil with a glr (gas-liquid ratio) of less than 3000 scf/stb (fig. 19d). further maturation would result in wet gas/condensate generation. regional implications the documentation of a rich gas/oil-prone upper jurassic succession in jameson land is obviously important for the assessment of the regional petroleum potential, including the north-east greenland continental shelf, the jan mayen area and perhaps also areas onshore greenland. in jameson land, the source-rock succession is intimately associated with potential reservoir rocks constituted by the various types of gravity-flow sands and injectites that make up more than 40% of the cored sequence, but the succession is near the surface or exposed over most of the area. however, the basin is gently tilted towards the south and seismic data indicate that the sedimentary succession continues under scoresby sund and the thick basaltic cover of the volquart boon kyst that faces jameson land from the south (larsen 1980). hence, a conceptual play could be developed, based on the intimate association of a rich gas/oil-prone source rock and interleaved units of sandstone showing good reservoir quality, sealed by overlying cretaceous–palaeogene mudstones and palaeogene volcanics. the cretaceous–palaeogene succession has been largely removed in jameson land during subsequent uplift (mathiesen et al. 1995, 2000; green & japsen 2018, this volume), but in areas south of 110110 jameson land and perhaps elsewhere in comparable settings it may have been preserved. such a play constitutes a variant of the synrift play demonstrated by the pil and bue discoveries on the halten terrace in the norwegian offshore. in the present case, stacked synrift sands act as reservoirs, separated by rich source rocks of the hareelv formation, equivalent to the melke and spekk formations in the norwegian example, which act as both source rocks and internal seals. conclusions the upper jurassic succession of jameson land includes a rich gas/oil-prone succession of marine black shales of oxfordian to early volgian age, deposited in deep shelfal oxygen-deficient environments that frequently received sandy gravity flows from shallower coastal areas. palaeogene basaltic intrusions of various sizes are commonly present. the deposits represent the local equivalents to the kimmeridge clay formation sensu lato known as the most important petroleum source-rock succession in the greater north atlantic region, and also in basins east of canada and in western siberia. outcrop samples have proven largely useless for assessment of petroleum potential because of intense weathering and alteration of the kerogen, probably due to sulfuric acid generated from the breakdown of abundant disseminated pyrite. the lollandselv-1, -2, falsterelv-1 and jyllandselv-1 fully cored boreholes together cover the lower oxfordian to lower kimmeridgian succession below, and overlapping with, the lower part of the blokelv-1 borehole succession. the deposits are characterised by high contents of terrestrial and inert organic matter but the reactive fraction of the kerogen has the capacity to generate both oil and gas. the sjællandselv-1, -2 and -3 fully cored boreholes partially overlap stratigraphically and together exhibit the uppermost kimmeridgian – lower volgian succession. the petroleum potential varies widely from zero to high for both oil and gas. although the succession seems to be developed in sandier facies and thus shows lower proportions of mudrocks, its character with respect to petroleum potential seems to match the corresponding part of the blokelv-1 borehole succession fairly well. due to its stratigraphic completeness and coverage from the middle oxfordian to the lower volgian – including kimmeridgian sea-level highstand deposits not represented elsewhere – the blokelv-1 fully cored borehole (td 233.8 m) serves as the reference for the entire succession. the succession penetrated by the blokelv-1 borehole consists predominantly of mudrocks (55%) and, based on petroleum potential/organic facies, the succession may be subdivided into two subunits, albeit the exact transition is obscured by the presence of a magmatic intrusion. the lower unit, extending from td to c. 110 m, is characterised by high levels of toc, moderate sulfur content, good s2-yields and hydrogen index values slightly below 300 on average, with a clear upwardsincreasing trend. these deposits accumulated during an overall rising sea level. the upper unit, showing fairly high levels of toc, predominantly oil-prone kerogen and very high levels of sulfur, and hi slightly above 300 on average, covers the interval from c. 90 m to 10 m with a maximum in most parameters in the interval 60–80 m. the sediments were deposited during a prominent sealevel highstand. the succession thus records changing depositional conditions in response to an overall rising sea level from td to c. 90 m, culminating in the interval 90–60 m, probably with a regressive tendency in the uppermost portion of the succession. the general depositional environment was marine and oxygen-restricted throughout the succession, but the kerogen facies changes from marine with a prominent terrestrial component towards increasingly marine compositions, culminating in the interval 90–60 m. this development is also recorded by a number of biological marker parameters as well as by the stable carbon isotopic ratio (δ13c). oil stains recorded in association with magmatic intrusions show biological marker fingerprints very similar to the shales adjacent to the intrusions, testifying to their local origin. hence, the intrusions have had a profound effect on the thermal maturity of adjacent rocks, but the effects are local and restricted to a fairly narrow zone below and above the intrusions. the characteristics of the oils fall well within the range of variation seen in oils generated from correlative successions elsewhere in the greater north atlantic region. a series of thermal maturity indicators (ro, tmax, biological markers) consistently show unexpected increasing depth-trends within only about 200 m of stratigraphic thickness. the thermal maturity of the blokelv-1 succession straddles the transition to the oil-generative window where a number of geochemical processes accelerate, and the trend in thermal maturity probably represents quenching of incipient petroleum generation by rapid neogene uplift after maximum depth of burial was reached in the palaeogene. calculation of the ‘ultimate expulsion potential’ (uep) by kinex™ modelling suggests that although the kerogen in general is classified as gas/oil-prone, ultimately about twice as much gas as oil would be generated from 111 the succession. the uep is unevenly distributed throughout the drilled succession with a maximum in the interval c. 170–130 m despite the lower average hi of this interval, when compared to the sea-level highstand succession above. a maximum estimate of the ultimate expulsion potential (uep) amounts to 61.4 mmboe/km2, which is in close agreement with published estimates of correlative successions in the greater north atlantic region. the documentation of a rich gas/oil-prone upper jurassic succession in jameson land is important for the assessment of the regional petroleum potential, including the north-east greenland continental shelf, the jan mayen area and perhaps also areas onshore greenland, where a new conceptual play may be defined that represents a variant of the ‘synrift sand play’ demonstrated by the pil and bue discoveries on the norwegian shelf. acknowledgments we wish to thank finn surlyk for helpful comments on an earlier version of this paper. the referees iain c. scotchman and erdem idiz are thanked for their highly useful and very constructive reviews. analytical work by ditte kiel-dühring and carsten guvad is gratefuly acknowledged, as is draftwork by jette halskov and stefan sølberg. references alsen, p. & piasecki, s. 2018: biostratigraphy of the hareelv formation (upper jurassic) in the blokelv-1 core, jameson land, central east greenland. in: ineson, j. & bojesen-koefoed, j.a. 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survey of denmark and greenland. bulletin 10, 65-70 the kenya coastline extends 600 km from the border of tanzania in the south to the border of somalia in the north (fig. 1). the kenyan coast features a diverse marine environment, including estuaries, mangroves, sea grass beds and intertidal reef platforms and coral reefs, which are vital for the reproduction of marine organisms. these coastal ecosystems are regarded as some of the most valuable in kenya but face serious threats from the ever increasing human pressure of tourism, industrial pollution, destructive fishing, mangrove logging and other unsustainable uses of marine resources. another serious threat is the maritime transportation activities along the coast and at the ports. it is estimated that at any given time more than 50 ships operate in the major shipping lanes off the kenyan coast, of which about nine are oil tankers with capacities ranging from 50 000 to 250 000 tonnes. furthermore, the harbour of mombasa serves as the major port for countries in east africa. in recognition of the risks posed by oil pollution the government of kenya and the commercial petroleum industry agreed to develop a national oil spill response contingency plan (nosrcp) with the purpose of enabling a speedy and effective response to any oil spill within the territorial waters of kenya. an important element of this plan was the mapping of the coastal resources and the development of an environmental sensitivity atlas showing the vulnerability of the coast to marine oil spills. in 2004, the government of kenya approached the united nations development program (undp) in kenya for financial support to develop an environmental sensitivity atlas. the project was approved and forwarded for funding by the danish consultancy trust fund administrated by united nations operational program (unops) in copenhagen. the project was announced in denmark, and the kensea group headed by the geological survey of denmark and greenland (geus) was awarded the contract. the project comprises four phases: (1) data compilation and development of the kensea database, (2) development of a coastal classification for kenya, (3) development of the sensitivity index jointly with a group of stakeholders, and (4) compilation of the kensea environmental sensitivity atlas (tychsen 2006). development of the kensea database in the 1990s the eastern africa coastal and marine resources database and atlas (unep 1998) was developed by the united nations environmental programme. this database contains a variety of datasets covering bioscience, geoscience and human use, which have been incorporated either directly or as guidelines for further investigations in the kensea database (kenseabase). however, the unep atlas has been produced at the scale of 1:250 000, implying a lower degree of spatial resolution than the scale of 1:50 000 requested for the kensea project. kensea – development of an environmental sensitivity atlas for coastal areas of kenya john tychsen, ole geertz-hansen and jesper kofoed © geus, 2006. geological survey of denmark and greenland bulletin 10, 65–68. available at: www.geus.dk/publications/bull 65 mombasa 50 km kenya africa 39°e 40°e 41°e 5°s 4°s 3°s 2°s kenya tanzania indian ocean somalia fig.3 fig. 1. the coastal area of kenya (red frame on index map) showing coverage of the kensea coastal sensitivity map (ruled area). green frame, location of maps shown in fig. 3. 66 following a kick-off seminar with all stakeholders, the team from the kensea group visited a large number of government agencies, ministries, private companies and national as well as international non-governmental organisations to collect additional and updated data. all collected data were digitised and included in the kenseabase, and preliminary maps were printed. based on these preliminary maps, a field programme for data verification and additional data collection was developed and carried out. the results were integrated into the kenseabase and logistic and topographic maps and coastal resource maps were produced. development of a coastal classification for kenya many attempts have been made worldwide to classify coastlines with respect to their vulnerability to oil pollution. these classifications are usually based on the geomorphology, the degree of exposure to wind and waves, and other relevant conditions. most of these classifications were inspired by gundlach & hayes (1978) and baker et al. (1994). the classification adopted for the kensea project is a further development of these classifications to meet the conditions in kenya. in addition to coastal geomorphology and degree of exposure, the ecological value and biodiversity of this particular stretch of coastline is taken into account, as well as the ability of the particular coastal types to facilitate selfcleaning and the conditions for removing a potential oil spill. the resulting classification comprises seven categories, each of which can be described in terms of its geomorphology, ecological value and vulnerability to an oil spill (tychsen 2006). rocky coast much of the kenyan coast is formed by 4–6 m high pleistocene coral limestone cliffs (fig. 2). at present-day sea level they are exposed to wave erosion, resulting in an irregular and rugged appearance. in the upper part of the intertidal zone biological activity is sparse. subtidal rocks and man-made hard-surface structures such as piers and wharfs may develop a richer flora and fauna, resembling the conditions found on reefs. the exposed cliffs are regarded as less sensitive to oil pollution than most other habitats because of the sparse biological activity, and because the exposure to waves makes the surface to some extent self-cleaning. sandy beaches and dunes two types of sandy shores are present along the kenyan coast: (1) gentle to steep sandy beaches without protection from a reef. the beach is often backed by one or a series of windblown sand dunes. the sand may be of terrestrial origin and supplied by the larger rivers. (2) gently sloping beaches sheltered behind a reef are common along the coast and the sand is often white calcareous sand of marine origin (coral sand). species diversity on sandy beaches is usually low. above the high-water line, only a few burrowing crabs and amphipods are usually found. the density and diversity of crabs, bivalves, polychaetes and other marine invertebrates increases in the intertidal zone, but remains low compared to most other habitats. fine-grained sandy beaches are less sensitive to oil pollution due to their relatively sparse biological activity. furthermore, they are relatively easy to clean since oil does not penetrate deep and can be removed either manually or by use of machinery. on the other hand, coarse sand or gravel is more sensitive as the oil can sink deep into it, and the oil may therefore be impossible to remove. coral reefs and reef flats most of the kenya coastline is fringed by a major barrier reef complex that includes the most diverse ecosystems in the marine environment. the up to 2 km wide shallow-water reef flats between the coastal cliffs and the reef crest comprise fossil reefs currently eroded by wave action. active reef growth occurs at the reef crest and on the slope facing the ocean where coral reef growth occurs to depths between 20 and 25 m. the main part of the reefs are subtidal, and are therefore sheltered from direct contact with a possible oil slick. however, reef crests are usually exposed at low tide and the intertidal corals will be killed immediately by contact with oil. the deeper parts of the reef may also be endangered as waves break on the crest and fine oil droplets become dispersed in the water column. recovery of damaged coral reefs may take several decades, and restoration techniques are usually not very successful. fig. 2. pleistocene coral limestone cliffs typical of the kenyan coast. rias rias are drowned river valleys or estuaries. the typical ria has a steep slope or a sheltered cliff, often with a narrow subtidal muddy beach with a few mangrove trees. the sheltered environment of the ria has only little self-cleaning capacity, although it is often possible to clean from the seaward side because of its narrow extent. river mouths and estuaries the rivers in the area form large, gently sloping floodplains with extensive estuarine zones characterised by fluctuating salinity. the mouths of smaller rivers are often hidden behind mangrove creeks. few plant species apart from mangroves are adapted to low or fluctuating salinities. the biodiversity is therefore low within the estuaries, although the density is usually very high due to the continuous supply of food and nutrient from the river. the high density of bivalves, snails and other benthic invertebrates usually attracts many birds. tidal currents can carry any oil pollution far into the estuary and thus into contact with a high density of food items. therefore the sensitivity is very high. flushing and self-cleaning is limited to the seasonal high flow situations. mangroves mangroves have a high productivity as they profit from nutrients from both land and sea, and mangrove detritus is often the main source of energy fuelling the estuarine food webs. mangroves are favoured by fine-grained nutrient-rich sediment, and are therefore often associated with estuaries and other freshwater outlets. while the diversity of mangrove tree species is limited, they create a multitude of niches suitable for a vast diversity of other organisms. oil can be acutely toxic to the mangrove as clogging of the aerial roots by oil may hinder proper ventilation 67 human use major town town settlement water intake historical/cultural site hotel ferry crossing biological resources turtle, high priority turtle, low priority forest bird wader, high sensitivity wader, moderate sensitivity wader, low sensitivity coral mangrove priority coral forest or thicket plantation papyrus swamp important bird area sand or mud seasonal swamp sensitivity index line level 1 (low) level 2 level 3 level 4 level 5 (high) coastal types mangrove sheltered rias tidal mud flats sheltered sand exposed sand rock coral environmental sensitivity map coastal resource map fish traps fish landing site air strip main road road track fig. 3. examples of a coastal resource map (left) and an environmental sensitivity map (right) from the kensea environmental sensitivity atlas (see fig. 1 for location). 68 of subsurface parts and lead to suffocation and stress, and eventually death. clean-up operations may prove extremely difficult due to the low energy environment of mangrove forests. intertidal mud flats sheltered mangroves in creeks or bays are often fringed by a broad intertidal mudflat. these are characterised by a high density of marine invertebrates such as mussels, snails and crustaceans. although diversity is usually low they are important feeding grounds for aquatic birds. the mudflats are sensitive to oil spills since they are difficult to clean mechanically without mixing the oil into the sediment. where mangrove forests are associated with mudflats the clean-up is even more difficult for both habitats. development of a sensitivity index the oil spill sensitivity ranking is designed to help decisionmakers to prioritise the available resources and to focus the emergency response on the most vulnerable areas, both (1) during the pre-spill planning, in order to prepare appropriate response strategies and (2) during an oil spill combat in order to plan and continuously optimise the response strategy under the given conditions and limitations (oil type, weather, equipment, crew, etc.). planning the clean-up operation and avoiding post-spill damage the goal of oil spill response is to minimise the overall impacts on natural and economic resources as well as cultural assets, but some aspects will be of greater concern than others. the sensitivity ranking for a given stretch of coastline therefore includes the actual sensitivity of the present resources or assets, and a more subtle evaluation of importance or value. the ranking should integrate a multitude of data such as geomorphological and geological properties, wave exposure, biological diversity and productivity, oil behaviour, ease of clean-up, human use and cultural assets. although these properties are not directly comparable or quantifiable the outcome should ideally be a simple statement, a numerical index value, or a colour code. the comparison and evaluation of incomparable properties is a matter of balancing often conflicting interests, and no perfect system can be devised. a complicated system is not necessarily better or more accurate than a simple system or qualified judgments. therefore, it is essential that the amount of detail matches the purpose, and leaves the decision-makers with some alternatives. the total sensitivity index developed for the kensea atlas embraces three main themes: (1) coastal type, (2) biological resource, and (3) human use. for each theme an index value that incorporates the index values for the various attributes encompassed by the particular theme has been allocated for a particular stretch of coastline. the index value for the individual attributes used in each of the themes and the formula for calculating the total index were developed during a seminar with participants from government agencies, ministries, private sector as well as national and international ngos. this joint development of the formula is crucial for the future use of the sensitivity index. the sensitivity index line has been reproduced on the environmental sensitivity maps (fig. 3). environmental sensitivity atlas the atlas covers the entire coastline and comprises three types of maps, each of which have been produced in 16 map sheets at scale 1:50 000 (fig 1; tychsen 2006). references baker, j.m., spalding, m.d., moore, j. & tortell, p. 1994:. sensitivity mapping for oil spill response. imo/ipieca report series 1, 24 pp. gundlach, e.r. & hayes, m.o. 1978: vulnerability of coastal environments to oil spill impacts. marine technology society journal 12, 18–27. tychsen, j. 2006 (ed.): kensea – environmental sensitivity atlas for coastal area of kenya, 76 pp. copenhagen: geological survey of denmark and greenland. unep 1998. eastern africa atlas of coastal resources, kenya, 119 pp. nairobi: unep. authors’ address j.t., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: jt@geus.dk o.g.-h., aquasim, slagslunde bygade 37, dk-3660 stenløse, denmark. j.k., geoquest, gyldenløvesgade 16, dk-1369 copenhagen k, denmark. geological survey of denmark and greenland bulletin 3, 151-171 151 total for climatic regions ...l: 1 (1.4%) the holocene skagen: 71 (28.7%) from the skagen boring, 71 holocene species have been recorded. the boreo-lusitanian group dominates with 54 species (76.1%) of the holocene mollusc species from the skagen well. one purely lusitanian species, vitreolina collensi, has been found, while all the other species occur in the boreal and to some extent the arctic. in this way the skagen well material resembles that of other regions like the vendsyssel, limfjord and north sea during the holocene. the skagen well material has all been recorded to certain stratigraphical levels, as seen in appendix 6, so the climatic indications through time appear, but the number of molluscan species is very low. in the preboreal/boreal, only three species have been recorded. this has been explained as a result of a deeper water where the echinoids dominate. higher up in the sequence, the number of molluscs increases – 23 species in the subboreal and 68 species in the subatlantic. through the chronostratigraphical levels, the climatic regions of the boreo-lusitanian from the dominating one, and the purely lusitanian vitreolina collensi as mentioned above occurs in the subboreal and subatlantic. however, as already stated, the development of the facies in the skagen well during the holocene does change the environment from the deeper-water facies with few molluscs through the bottom community with turritella communis into the prolific shallow-water community. in this way the youngest part covering the subboreal–subatlantic is also by far the part with the highest species diversity. the environmental changes within the seven regions through the late quaternary evaluated by the molluscan communities met with in the seven stages the seven chronological stages which have been described according to their climatic affinities are seen in fig. 102 and fig. 103 covering the eemian, the early/ middleweichselian, thelateweichselian, thepreboreal/ boreal, the atlantic, the subboreal, and the subatlantic. in this way the climatic cycle during the late quaternary is demonstrated on the basis of marine mollusc species which indicate that the eemian has by far the highest amount of the more temperate species, while the holocene reached its maximum during the atlantic, although only slightly more than the other stages within the holocene, as already commented upon in the previous chapter. it is generally accepted that the eemian summer temperatures were higher – about 2°c above the present. with glaciers smaller than the present day, this means that the sea level was 4–6 m higher than today (andersen & borns 1994, pp. 44–49). and as pointed out by donner (1995, p. 39): “the submergence was clearly greater after the saalian glaciation than after the weichselian and possibly after the older glaciations”. donner sees this in northern europe as “a result of a comparatively great downwarping of the earth’s crust during the extensive saalian glaciation”. the rebound since the last glaciation has come to an end within the danish area (petersen 1985c, 1991b). this means that the eemian deposits, when found in denmark in nonglacio-dislocated state, can be regarded as being in the original position related to sea level, although there might be some movements in relation to neo-tectonic activities, as mentioned earlier. in the light of the observations mentioned above, the seven regions will be discussed according to the environmental characteristics such as the climatic affinities for the molluscs recorded in appendix 6 for each region, as appearing in fig. 103. however, for the holocene still as many as 130 species including the recent ones (95) not found as subfossil have not been dated to give their first appearance, see fig. 102: unknown arrival in holocene. at the end of each of the seven stages the molluscan communities sensu c.g.j. petersen will be presented in tables 3–9. geus bulletin no 3.pmd 09-07-2004, 09:10151 152 eemian stage 130 000 – 115 000 b.p. the bælt sea, region 1 appendix 6 and fig. 103 already forchhammer (1842, p. 64) designated cyprina (arctica) islandica to be the characteristic bivalve of the bælt sea eemian, as known to the present-day geologists. furthermore, forchhammer points out that the characteristic bivalve, cyprina islandica, occurs every where in large quantities, but always in crust specimens. however, all the shell fragments occurring together show that the specimen has been broken after deposition in the clay, most probably by the cataclysms which have given the beds their tilt. johnstrup (1882a, p. 55) points to the indications of the molluscs as being a deposition of a shallow-water sea and also mentioned the mytilus beds. johnstrup points out (1882a, p. 56) that the dislocated floes – as already noticed by forchhammer – have the original succession within each floe, saying that the cyprina clay and the mytilus beds have not been disconnected during the dislocations. later investigations by nordmann (harder 1900; nordmann 1908, 1913) demonstrated that the venus aurea as observed by johnstrup (1882a, p. 66) could be regarded in parts as the no longer living tapes aureus gm. var eemiensis nordmann or tapes senescens doederlein; in this book paphia aurea senescens. the tapes species do represent shallow-water environments (see the chapter on the molluscan species), and therefore the whole of the bælt sea region can be characterised by the three bivalves mentioned above, from the eulittoral to the infralittoral shallow-water zones: mytilus, tapes and cyprina. among the three species mentioned, the tapes species (paphia aurea senescens) also remains as the only subfossil bivalve from denmark which can be regarded as an index fossil from the marine eemian. the stratigraphical position of the marine eemian is according to jessen & milthers (1928, p. 179) contemporaneous with the mixed oak forest zone and the carpinus zone in the interglacial bogs; furthermore, jessen & milthers conclude (1928, p. 341) that the climate of jylland and nw germany in that part of the interglacial period which answers to zone f [culmination of the curves for mixed oak forest] was no less atlantic in character than (100%) (100%) (100%) (100%) (100%) (100%) (100%) 3 11 48 48 16 53 3 290 14 140 36 41 27 93 77 72 486 1 2 6 5 10 19 14 1 1 2 1 11 7 11 2 6 5 6 3 2 4 2 3 2 15 1 5 3 14 7 8 1 1 1 90 21 70 56 53 10 3 1 (3%) (5%) (17%) (12%) (7%) (53%) (34%) (4%) (1%) (3%) (1%) (8%) (19%) (27%) (7%) (6%) (6%) (8%) (2%) (6%) (10%) (2%) (4%) (3%) (11%) (3%) (12%) (11%) (15%) (9%) (11%) (1%) (1%) (1%) (64%) (78%) (75%) (73%) (74%) (7%) (4%) (1%) frequency: number of molluscs column per cent, % climatic affinity eemian early/middle late weichselianweichselian preboreal/ boreal atlantic subboreal subatlantic total number of species unknown arrival in holocene arctic arctic, subarctic arctic, subartic, boreal, lusitanian arctic, subarctic, boreal subarctic, boreal subarctic, boreal, lusitanian boreal boreal, lusitanian lusitanian total number of species column per cent, % 35 s ub fo ss il sp ec ie s 95 r ec en t sp ec ie s pleistocene holocene time: the seven stages fig. 102. the seven stages from the eemian through the weichselian – comprising the early/middle weichselian and late weichselian – to the end of the holocene. preboreal–boreal, atlantic, subboreal, and subatlantic are here shown according to their climatic affinities based on the molluscan record. geus bulletin no 3.pmd 09-07-2004, 09:10152 153 (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (2%) (3%) (10%) (73%) (12%) (2%) (10%) (7%) (9%) (28%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (5%) (5%) (5%) (21%) (63%) (2%) (2%) (5%) (5%) (3%) (3%) (9%) (72%) (16%) (2%) (4%) (5%) (20%) (1%) (5%) (1%) (11%) (73%) (9%) (2%) (8%) (5%) (12%) (14%) (32%) (16%) (13%) (2%) (11%) (2%) (56%) (16%) (11%) (5%) (7%) (33%) (7%) (7%) (29%) (14%) (43%) (7%) (2%) (6%) (2%) (1%) (4%) (14%) (14%) (29%) (43%) (17%) (8%) (4%) (5%) (3%) (17%) (53%) (19%) (6%) (3%) (17%) (50%) (33%) (11%) (10%) (1%) (25%) (75%) (17%) (5%) (6%) (11%) (37%) (23%) (9%) (14%) (33%) (33%) (23%) (13%) (15%) (6%) (17%) (8%) (5%) (5%) (5%) (4%) (8%) (12%) (77%) (2%) (3%) (4%) (4%) (33%) (67%) (1%) (0%) (6%) (3%) (16%) (74%) (3%) (5%) (6%) (5%) (13%) (87%) (2%) (3%) (22%) (78%) (6%) (4%) (1%) (8%) (3%) (17%) (71%) (2%) (9%) (10%) (15%) (12%) (7%) (7%) (85%) (3%) (2%) (5%) (50%) (50%) (5%) (0%) (100%) (2%) (50%) (50%) (1%) (0%) (6%) (6%) (6%) (78%) (6%) (4%) (3%) (2%) (6%) (8%) (16%) (11%) (5%) (5%) (63%) (5%) (10%) (1%) (33%) (3%) (33%) (67%) (1%) (0%) (4%) (4%) (9%) (78%) (4%) (2%) (5%) (2%) (4%) (4%) (100%) (0%) (13%) (25%) (63%) (2%) (2%) (1%) (1%) (7%) (3%) (10%) (1%) (75%) (1%) (2%) (8%) (10%) (8%) (33%) (11%) (4%) (11%) (11%) (33%) (33%) (11%) 59 19 32 91 55 14 7 36 4 35 9 26 3 31 15 23 77 27 2 8 2 36 19 3 23 1 8 68 7336 12 57 64 20 84 3 462 25 1 1 2 18 1 1 1 2 5 1 5 2 7 1 51 1 8 1 1 2 2 2 28 2 3 2 1 1 12 1 2 1 2 2 1 5 23 2 13 5 18 1 6 2 13 55 2 2 23 1 1 1 1 2 3 1 6 19 7 2 1 1 3 2 4 13 8 3 5 1 1 3 3 1 1 2 3 20 1 2 6 43 7 1 1 1 4 12 1 3 23 5 1 5 1 10 66 8 9 7 1 6 1 31 1 4 2 6 1 age ep oc h pl ei st oc en e h ol oc en e region 1. bælt sea 2. baltic sea 3. kattegat 5. north sea 6. vendsyssel 7. skagen 3. kattegat 6. vendsyssel 7. skagen 6. vendsyssel 7. skagen 5. north sea 7. skagen 1. bælt sea 2. baltic sea 3. kattegat 4. limfjorden 5. north sea 6. vendsyssel 7. skagen 3. kattegat 4. limfjorden 5. north sea 6. vendsyssel 7. skagen 4. limfjorden 5. north sea 7. skagen eemian early/middle weichselian late weichselian preboreal/ boreal atlantic subboreal subatlantic total of species in number row per cent, % column per cent, % total of species: in number row per cent, % column per cent, % arctic a, s boreal b, l lusitanian a, s, b a: arctic, s: subarctic, b: boreal, l: lusitanian a, s, b, l s, b s, b, l climatic affinity fig. 103. climatic characterisation (affinity) of mollusc assemblages for each region over the seven stages (ages) since the eemian. geus bulletin no 3.pmd 09-07-2004, 09:10153 154 the climate of the litorina period in postglacial time. also s.t. andersen (1965, pp. 499–500) points to the eemian as having an oceanic and warm climate. in his diagram from hollerup, zone 5 represents the climax forest. jessen & milthers (1928, p. 179) bring forth the view of nordmann saying that the lusitanian mollusc fauna was moving into the baltic basin by way of the sounds that cut diagonally through the southern portion of the cimbrian peninsula. however, this idea is, as discussed earlier, not accepted by the present author, where a connection to the west is advocated to be over the kattegat–vendsyssel–skagen regions. an eastern open connection to the white sea over finland and russia will be discussed in connection with the baltic region. the baltic, region 2 appendix 6 and fig. 103 in the baltic sea part of denmark the number of recorded marine molluscs from the eemian has fallen to 19 species compared to the bælt sea region, and no purely lusitanian species occur. however, as pointed out earlier, this change in the climatic situation regarding the baltic fauna can be seen as a consequence of the fact that here only the deeper-water assemblage characterised by turritella communis occurs, although one of the species known from the characteristic part of the bælt sea fauna is recorded, i.e. arctica islandica. however, the tapes species are not met with in this region. the deeper-water environment is therefore well characterised by the turritella community. although the decline in number of species very much resembles the present-day situation between the bælt sea and the baltic as mentioned earlier, five species in the eemian fauna, including turritella communis, show a salinity above the present conditions in the baltic. from the study of diatoms at ollala in eastern fennoscandia, forsström et al. (1988, p. 322) write: “this mixture of warm and cold indicators probably means that the eemian sea in the baltic basin had a connection both to the north sea in the west and to the arcticocean via thewhite seabasin in thenortheast”. only a few works from eastern fennoscandia have been based on molluscan studies. however, among the papers by zans (1936), sokolova et al. (1972), and gross (1967), gross mentions the following molluscs: portlandia arctica, clinocardium ciliatum, heteranomia squamula, macoma calcarea, littorina littorea, and cerastoderma edule. here the three last mentioned species occur in the eemian from the danish baltic and clinocardium ciliatum from the vendsyssel region, although here in the upper turritella terebra zone correlated to the early weichselian, as discussed later. portlandia arctica has not been recorded from the danish eemian, although it occurs in the beds below the eemian in the anholt boring (the kattegat region), where seidenkrantz (1993, p. 284) also has demonstrated foraminiferal zones a–d with arctic species. gross (1967, p. 118) regards the arctic and arctic– boreal molluscs in the eemian clay as: “relikte aus der portlandia-transgression des dnepr ii-spätglazials, die nach dem pollen-profil und -diagramm der eem-transgression voranging”. the older correlation of the so-called weissmeer transgression by zans (1936, table 1) contains further details on the molluscs upon which the correlation has been based, and it also includes the danish area, mentioning ‘dänische inseln’ after ødum (1933) and ‘skaerumhede’ after jessen et al. (1910). however, the occurrence of the high-arctic portlandia arctica should be placed in the late saalian, as seen in the kattegat region mentioned above, while the arctic–boreal species clinocardium ciliatum could be taken as a relict in the skærumhede sequence from the late saalian environment within the danish area or introduced by the cooling in the early weichselian. the faunal development has been worked out in more detail between the eastern fennoscandia and the danish area (funder et al. 2002). the connection to the arctic over the white sea during the eemian seems to be well established, but only for a shorter time, 1000–2000 years of the more than 10000 years that the eemian sea existed in the baltic region (funder 2000, p. 68). the kattegat, region 3 appendix 6 and fig. 103 from the kattegat region, 32 species have been recorded, with a high amount of boreo-lusitanian species (72%). five species are lusitanian, among which are found the characteristic eemian species of the shallow-water environment, including the tapes species. however, also the deeper-water environment is represented by the turritella community in this region. geus bulletin no 3.pmd 09-07-2004, 09:10154 155 on anholt, the turritella community occurs at a depth of around 70 m b.s.l., and the tapes fauna in the isefjord area at ejby at a depth of around 10 m a.s.l. the latter is considered to be in situ (madsen 1968). it is tempting to regard the two localities as being at about their original elevation in relation to an eemian sea level some what higher than the recent one, since the glacio-isostatic rebound had expired (petersen 1991b). on the basis of the scattered eemian localities of which some are floes in the weichselian glacial deposits, the maximum extent of the eemian sea cannot be given. however, both the shallow-water environment characterised by the tapes species and the deeper water by the turritella species have been demonstrated. in this way both of the characteristic marine environments from the bælt sea and the baltic respectively are represented in the kattegat region. the north sea, region 5 appendix 6 and fig. 103 the largest amount of mollusc species within the eemian have been recorded from the north sea region, or to be more precise from the coastal region of the north sea. in the danish part of the north sea, many studies on microfossils from the oil and gas fields have demonstrated eemian deposits in the central north sea, but their macrofossils have not been studied (knudsen 1985a, 1986). however, the large amount (91) of molluscs from the coastal region fall into different facies, as seen in the previous regions when a much lower number was looked at. gripp (1964) uses the ‘senescens sand’ and ‘turritella ton’ to give his idea of the marine order of the strata. however, in this context, working with regions and not with localities, it should be emphasised that the development of different facies most probably happened in parallel. gripp (1964, p 223) expresses this himself in saying: “tapes-sand und turritellen-ton sind die beiden facies, die während des ansteigs des meeresspiegels entstanden”. as seen from the species found in the danish north sea coastal region, we do find the turritella species and tapes species, but also the occurrences of donax vittatus should be mentioned as a facies indicator, characterising the high-energy coastal environment from this area facing the eemian north sea. from recent studies on foraminifera in northern germany at the kiel canal, knudsen (1986) shows that the marine transgression took place in the warm part of the eemian, and hinsch (1985) in his mollusc study from the same area revealed three mollusc communities characterising the shallow-water environment, with such genera as mytilus-cerastoderma, acanthocardiavenerupis and bittium-varicorbula. however, the old material from many localities in the danish north sea region cannot be worked out to such detail, although all the marine mollusc species mentioned by hinsch (1985) have been recorded from the danish eemian north sea region. when the molluscan fauna in the danish eemian north sea region is compared with the eemian on the west coast of norway as described by mangerud et al. (1981), 20 molluscan species out of the 35 species recorded from the fjøsangerian are known from the north sea region and 7 species from other eemian regions in denmark. here macoma calcarea and nuculana pernula belonging to the arctic–boreal group occur in the vendsyssel and skagen regions, and in the baltic region macoma calcarea representing deeper water during the eemian, while chlamys islandicus, which does not occur in the danish eemian deposits, is found in the weichselian recorded from the vendsyssel region. this means that the fjøsangerian can be regarded as slightly cooler than the danish north sea eemian deposits. another marine eemian deposit in norway described by andersen et al. (1983) at bø on karmøy (sw norway) revealed 25 molluscan species from the avaldsnes interglacial described in detail by sejrup (1987). here as many as 20 species are in common with those in the danish north sea region, and one, hinia incrassata, has been recorded from the eemian in the vendsyssel region. four species have not been found in the danish eemian. these are the arctic–lusitanian and arctic– boreal species puncturella noachina and boreotrophon clathratus respectively, the latter occurring in the late weichselian deposits in the vendsyssel region. this is much in line with the observations from the fjøsangerian deposits. the four non-occurrences in the danish eemian among the boreo-lusitanian species pecten maximum and lucinoma borealis, the latter being common at both norwegian localities, are difficult to explain. among the 8 purely lusitanian species recorded from the north sea region, only plagiocardium papillosum occurs in the norwegian eemian at fjøsanger, which again points to a slightly cooler position for the norwegian localities. geus bulletin no 3.pmd 09-07-2004, 09:10155 156 the vendsyssel region, region 6 appendix 6 and fig. 103 the eemian mollusc fauna from vendsyssel comprises 55 species with no purely lusitanian climatic affinity, while quite a few (nine species~ 16%) are found in the arctic and the boreal zones. the stratigraphical position has been well elucidated through foraminiferal investigations (knudsen & lykke-andersen 1982; knudsen 1984, 1985b, 1992; lykke-andersen 1987). the study by lykke-andersen (1987, fig. 5) also involves the molluscs, and references are made to the zones established on the basis of macrofossils. according to the foraminiferal studies, the transition to the early weichselian takes place around 120 m b.s.l. in the skærumhede i sequence (jessen et al. 1910), which is about 140 m b.s. this means that the upper part of the turritella terebra zone falls within the early weichselian. the two arctic to high boreal species serripes groenlandicus and clinocardium ciliatum at depths of 132 m and 127 m b.s. respectively are discussed by nordmann (jessen et al. 1910, pp. 124–128), and the climatic indications from turritella communis mean that the assemblage existed at the transition between the high and middle boreal. in the paper by knudsen (1992), it is said that an abrupt faunal change at the eemian–weichselian boundary reflects a drop in water depth of at least 50 m and a subsequent drop in temperature of several degrees. the drop in temperature might well be reflected in the mollusc fauna by the occurrences of the two bivalves mentioned above, and for the drop in sea level it is tempting to recall the observed occurrence of the eulittoral mytilus edulis at a depth of 135 m b.s.l. in the skærumhede i boring and up to the abra nitida zone, which forms the transition to the arctic turritella erosa community as mentioned earlier. therefore, within the vendsyssel area the eemian (isotopic stage 5e) is represented by a turritella community that continues into the beds representing the isotopic stages 5d–a (knudsen 1992, fig. 4). the hordalandian stage in western norway contains serripes groenlandicus and clinocardium ciliatum species and is referred to the early weichselian (mangerud et al. 1981). arctic conditions first occurred in the macrofossil zones turritella erosa, balanus crenatus and macoma calcarea (bahnson et al. 1974), which cover the portlandia arctica zone sensu nordmann (jessen et al. 1910, fig. 8). the skagen region, region 7 appendix 6 and fig. 103 the 14 eemian molluscs found in the skagen boring are the lowest number recorded within the eemian sites. however, the finds are a clear omen of the deeperwater environment not encountered earlier in denmark on the basis of molluscs. the recorded molluscs point to an environment like the deeper part of the skagerrak today, with a community such as the amphilepis norvegica/pecten vitreus, where the latter (delectopecten vitreus) occurs in the skagen well, as mentioned earlier. the boundary to the overlying arctic deposits characterised by the occurrences of portlandia arctica is sharp and coincides with a sedimentological change to a diamicton with dropstones in the arctic part, as found in the skærumhede sequence within the turritella erosa zone (bahnson et al. 1974). therefore, in the skagen well no transition zone from substage 5e to 5d–a can be demonstrated in the molluscan faunas. the recorded eemian communities and/or characteristic molluscan species for six regions with eemian marine deposits are given in table 3. 1. bælt sea mytilus edulis littoral tapes spp. shallow arctica islandica deeper 2. baltic littoral shallow turritella t. communis deeper 3. kattegat mytilus edulis littoral tapes spp. shallow turritella t. communis deeper 5. north sea donax vittatus littoral tapes spp. shallow turritella t. communis deeper 6. vendsyssel littoral shallow turritella t. communis ~ 100 m 7. skagen littoral shallow amphilopsis/pecten delectopecten vitreus > 100 m region community species depth table 3. eemian communities and/or characteristic molluscan species geus bulletin no 3.pmd 09-07-2004, 09:10156 157 early/middle weichselian stage 115 000 – 25 000 b.p. the kattegat, region 3 appendix 6 and fig. 103 although the number of mollusc species recorded from the kattegat region during the early/middle weichselian is low – seven species – the climatic indications for the arctic environment are clear, considering that all species canbe found in thehigh arctic, and that one species, portlandia arctica, is high arctic par excellence, and macoma calcarea indicates shallow water. the stratigraphical position of the arctic macoma community found in the kattegat region – holmstrup on sjælland – has been determined by foraminiferal correlation and aminostratigraphical investigations, as mentioned earlier (petersen & buch 1974; miller & mangerud 1985). recently, the foraminiferal studies of the quaternary sequence in the anholt boring have demonstrated a middle weichselian deposit at a depth of about 50 m b.s.l. (seidenkrantz 1993). considering the information given by knudsen (1992) on a drop of sea level of around 50 m during the transition from the eemian to the weichselian, the middle weichselian beds in the cored section on anholt may represent rather shallow-water deposits. this is in accordance with the occurrence of the arctic macoma community. the vendsyssel region, region 6 appendix 6 and fig. 103 foraminiferal studies by lykke-andersen (1987) indicate that the early weichselian beds are represented by the upper turritella terebra zone, and that the deeper-water temperate turritella community continued into the first part of the weichselian with the following abra nitida zone as a transition to the arctic deeper-water turritella community (turritella erosa). together with the arctic and arctic–subarctic species (seven in number, forming nearly 20% of the mollusc species), the sedimentological data show the occurrences of ice-rafted material recorded both from the skærumhede i and the skærumhede ii borings (jessen et al. 1910, p. 76; bahnson et al. 1974, figs 3, 4, 7) reflecting arctic conditions. it has been argued by nordmann (jessen et al. 1910) that species within the genera mytilus, cyprina, zirphaea, nassa, and bittium must be regarded as allochthonous and older elements. however, they could also be regarded as stray finds from contemporary shallow-water to littoral deposits occurring within times of higher temperatures in the near-shore areas, similar to the near-shore fauna of middle weichselian age – the bø interstadial (40–64 ka) with gibbula cineraria and mytilus edulis, the latter occurring frequently (sejrup 1987). during the younger part of the marine middle weichselian – around 32 000 b.p. – when the shallow-water arctic macoma community was established, no mytilus edulis or bittium reticulatum have been recorded. the development of the bottom communities within the older weichselian sequence is therefore given by the transition from the turritella communities in deeper water to the arctic macoma community in shallow water. the skagen region, region 7 appendix 6 and fig. 103 very few molluscs have been found in the skagen boring of older weichselian age. the four species are all arctic, and the occurrence of portlandia arctica shows that high arctic conditions have prevailed and ice-rafted material occurs. there is no indication of near-shore fauna as recorded from the vendsyssel region. from this, it might be concluded that the deposition of these beds took place in the first part of the middle weichselian, contemporaneous with the deposition of the turritella erosa beds of the skærumhede sequence, but at a water depth of more than 100 m, as demonstrated earlier. from the few finds, it is not safe to point to a certain community on the basis of molluscs. however, the community in deeper arctic waters is described by other animals than molluscs, i.e. the ascidia–spongia epifauna, and at depths exceeding 200 m by gorgonocephalus species. from the estimate on water depth taken in comparison with the early part of the arctic sequence in the vendsyssel region, the palaeodepth must have been well above 100 m. the recorded early/middle weichselian communities and/or characteristic molluscan species are given in table 4. geus bulletin no 3.pmd 09-07-2004, 09:10157 158 late weichselian stage 25 000 – 10 000 b.p. the vendsyssel region, region 6 appendix 6 and fig. 103 the 35 mollusc species found in the vendsyssel region are in number very close to the number of species encountered in the vendsyssel region during the early/ middle weichselian (36 species). however, as seen in fig. 103, the percent of molluscs with a wide range and only connected to the subarctic and southwards is higher in the late weichselian (46%) compared to the early/middle weichselian (28%). this can be explained by the way the development in the two seas before and after the main glaciation, the older and younger yoldia sea respectively took place. the deposits from the older yoldia sea reflect the transition from deeper arctic to shallow-water arctic communities, the turritella and macoma communities respectively. the late weichselian beds within the shallow-water environment show a development from the arctic macoma community to the boreo-arctic mytilus-zirphaea community after 13 000 b.p., with a deeper-water community characterised by the portlandia arctica species, as outlined by petersen (1984), which could be part of the deeper macoma community – the so-called ophiocten zone. these observations form the background for the earlier given explanation of the occurrences of boreal shallow-water species such as mytilus edulis in the deeper-water arctic community in the older weichselian deposits from the vendsyssel region. the occurrences of the mytilus edulis species in the late weichselian deposits in large quantities are described by jessen (1899). the dates of the earliest occurrences of zirphaea and mytilus go back to 12 770 and 12 520 b.p. 14c years respectively. all the 30 14c dates forming the base for the evaluation of the late weichselian sea levels and occurrences of fauna communities as figured in petersen (1984, fig. 1) have been listed by petersen & rasmussen (1995a, table 1). it appears that the dates older than 13 000 b.p. 14c years all come from hiatella arctica and macoma calcarea (only one date) going as far back as 14 650 ± 190 b.p. 14c age. considering the mollusc species inhabiting the waters “of the swedish west coast shortly after deglaciation” (fredén 1986, p. 55), one finds also chlamys islandica and mytilus edulis shortly after 13 000. the latter within the time span of its first dated occurrence in the younger yoldia sea deposits in denmark. as to the deeper-water deposits also around 13 000, one can take the sample from the dybvad clay pit (fig. 1) dated to 13 010 ± 190 b.p. 14c which contained the following molluscs: hiatella arctica, mya truncata, macoma torelli, portlandia arctica, buccinum groenlandicum and cylichna occulta (petersen 1984). it is seen that still after the immigration of boreo-arctic fauna to the shallow-water environment, showing an amelioration in climate, the arctic community persisted in the deeper water as shown by the dating from the bindslev clay pit giving 12 650 ± 180 14c age b.p., with species such as macoma calcarea and portlandia arctica. the same situation can be estimated from the investigation of the pleistocene/holocene boundary in south-western sweden (the moltemyr core) where “zone z comprising the samples from 560 cm to 650 cm, is characterised by portlandia arctica and nucula tenuis (nuculoma tenuis, here taken as a species with a wide climatic range), and by the absence of many of the species of the overlaying zone (such as mytilus edulis) … the water depth during deposition of zone z was greater than during any of the other zones (above) probably more than 20 m” (feyling-hanssen 1982, p. 128). regarding the climatic indication of portlandia arctica, feyling-hanssen (1982, p. 131) quotes andersen (1975, p. 54) saying: “evidently, portlandia arctica lived near the ice fronts [in southern norway] also during older, glacial phases, but it seems to have disappeared from our coasts shortly after the ra event, probably due to a warming of the sea”. 3. kattegat littoral arctic macoma m. calcarea shallow deeper 6.vendsyssel littoral arctic macoma m. calcarea shallow arctic turritella t. erosa ~ 90 m turritella t. communis ~ 90 m 7. skagen littoral shallow ascidia–spongia 100–200 m region community species depth table 4. early/middle weichselian communities and/or characteristic molluscan species geus bulletin no 3.pmd 09-07-2004, 09:10158 159 according to sørensen (1979), the disappearance of portlandia arctica from the oslo fjord area occurred somewhat before 10 000 b.p. the younger dryas marine deposits have not been demonstrated in the vendsyssel region but are recorded from the skagen region, which will be discussed next. the skagen region, region 7 appendix 6 and fig. 103 the nine species recorded from the skagen well indicate an arctic deeper-water community very much like the arca-astarte crenata community (ockelmann 1958). there are no finds of species which could be referred to the more shallow-water environment as seen in the case of the vendsyssel region to the south both during the early/middle weichselian and the late weichselian. through most of the history of the cored section of pleistocene age, the skagen well has revealed mollusc assemblages from a deeper-water environment. also the transition to the holocene takes place in deeper water. the purely arctic species portlandia arctica together with bathyarca glacialis is found right up to the strata dated to around 10 000 b.p. forming the pleistocene– holocene boundary in the skagen well. however, in the description of the shell fauna of the marine clays in the oslo fjord region, brögger (1900, p. 685) states: “portlandia arctica is never found in the arca clay”. later datings of the ‘middle arca clay’ and the ‘younger arca clay’ given by b.g. andersen(1965,p.118) yielded early preboreal ages. in the younger arca clay from norway, species such as mytilus edulis, zirfaea crispata, and macoma balthica (b.g. andersen 1965, table 2) are also found, which characterise the shallowwater deposits in vendsyssel after 13 000 b.p. 14c age. the recorded recent occurrence of bathyarca glacialis from southern iceland implicates extension into the high-boreal region, although the main extension is in the arctic. the norwegian records of bathyarca glacialis come from a more shallow-water environment, as seen from the occurrences of the three shallowwater species mentioned above. the recorded late weichselian communities and/or characteristic molluscan species are given in table 5. the preboreal–boreal stage 10 000 – 8000 14c years b.p. the north sea, region 5 appendix 6 and fig. 103 mollusc faunas from the late weichselian have here been recorded only from the vendsyssel and skagen regions. as seen in fig. 103, the early part of the holocene, the preboreal and boreal, have a record of 26 species, with as many as 77% (20 species) boreo-lusitanian. this is in contrast to the records from the late weichselian, when the arctic–boreal elements dominated, with 54% in the vendsyssel region and 55% in the skagen well, the latter with only a few species and representing a deeper-water environment. the preboreal–boreal north sea faunas contain eulittoral as well as shallow-water species. mytilus edulis, littorina littorea and cerastoderma edule characterise the littoral zone and macoma balthica the shallow-water zone. using the characteristic species from the c.g.j. petersen community concept, the oldest recorded faunal communities from the north sea might be the mytilus epifauna community with littorina littorea, and the macoma infauna community with cerastoderma edule. also the abra community on mixed bottoms with phaxas pellucidus, corbula gibba and mya truncata might be reflected in the recorded species. the skagen region, region 7 appendix 6 and fig. 103 as discussed earlier, the environment of the earliest holocene, the preboreal and boreal, can be referred to the maldane-ophiura sarsi community. this deep6. vendsyssel mytilus/zirphaea z. crispata littoral arctic macoma m. calcarea shallow ophiocten zone portlandia arctica deeper 7. skagen littoral shallow arca-astarte bathyarca glacialis deeper region community species depth table 5. late weichselian communities and/or characteristic molluscan species geus bulletin no 3.pmd 09-07-2004, 09:10159 160 water community observed in the skagen region compared to the shallow-water communities recorded from the north sea coastal region once more demonstrates the unique position of the deeper-water communities observed in the skagen well material compared to the other regions in denmark through the late quaternary. the recordedpreboreal–boreal communities and/or characteristic molluscan species are given in table 6. the atlantic stage 8000–5000 14c years b.p. the bælt sea, region 1 appendix 6 and fig. 103 out of the 47 species known from the bælt sea region, 31 have been recorded from the atlantic (fig. 103). the littoral zone with mytilus edulis, littorina littorea, littorina saxatilis, cerastoderma edule, and macoma balthica from the shallow-water zone, is recorded; furthermore, the abra alba community together with corbula gibba. these faunal elements reveal the mytilus epifauna community with ostrea edulis, which no longer occurs in this area. in the infralittoral zone is the macoma infauna community where such species as paphia aurea, tapes decussatus, and venerupis pullastra have been found, and finally the abra community with corbula gibba. also the epifauna on the vegetation is reflected in rissoa albella, r. membranacea, and r. inconspicua and other gastropods. also the bivalve parvicardium exiguum is associated with the vegetation. the communities mentioned are still to be found in the bælt sea region, whereas paphia, tapes, venerupis and ostrea are no longer found in this region. the baltic, region 2 appendix 6 and fig. 103 fifteen mollusc species out of the 19 recorded from the baltic region during the holocene can be referred to the atlantic (fig. 103). theoccurrencesofmytilusedulisandlittorina littorea are referred to the mytilus epifauna community in the littoral zone, while species, as macoma balthica, cerastoderma edule, and scrobicularia plana represent the macoma infauna community in shallow water. the occurrences of both littorina littorea and scrobicularia plana are characteristic for the atlantic in the baltic and are now absent. spärck (1950) points to the wider extent of scrobicularia plana in the stone age as a consequence of warmer water in those days; however, in the present work the higher salinity is preferred as an explanation, as mentioned earlier. this is supported by the occurrences of gastropods like bittium reticulatum, rissoa albella, and aporrhais pespelicani, species recorded from other regions today with higher salinity. it should be noticed that scrobicularia plana and littorina species have been demonstrated as far north in the baltic as estonia (kessel & raukas 1979, fig. 9), although only with a low percentage but persisting into the subboreal. the unexpected find of ostrea edulis from estonia has later been re-evaluated as transported there by some seamen and thrown then overboard (in a letter from prof. a. raukas, may 1995). according to nordmann (1903b, 1906), madsen (1944), and spärck (1942, fig. 21), the southernmost finds of subfossil oysters are the bælt sea and øresund off landskrona. the kattegat, region 3 appendix 6 and fig. 103 from the kattegat region, only half of the recorded species have been dated so as to give a first appearance date. this amounts to 23 species from the atlantic (fig. 103). all of the dated species come from geologically mapped areas and not from the kattegat proper. therefore the observed species all come from shallow-water environments, excluding the deeper-water environment recorded from foraminifera (christiansen et al. 1993; seidenkrantz & knudsen 1993). 5. north sea mytilus m. edulis littoral macoma m. balthica shallow abra a. alba deeper 7. skagen littoral shallow maldana/ophiura > 150 m region community species depth table 6. preboreal–boreal communities and/or characteristic molluscan species geus bulletin no 3.pmd 09-07-2004, 09:10160 161 from the listed species dated to the atlantic the mytilus epifauna community with littorina littorea and the macoma infauna community with cerastoderma edule and tapes decussatus can be pointed out. furthermore, the abra community with corbula gibba, which is common in present-day inner danish waters (thorson 1950), is present. the limfjord, region 4 appendix 6 and fig. 103 this region has the highest number of recorded mollusc species from the holocene, viz. 147 species, and 77 have been dated to the atlantic (fig. 103). petersen (1918, pp. 22–36) described the communities in the limfjord region covering the macoma balthica, the venus and the abra communities. also an area with mya truncata is mentioned, forming a transition zone between the macoma and the abra communities. in patches the epifauna elements such as mytilus edulis and modiola modiolus are found. from the zostera vegetation, the rissoa and bittium species are mentioned. all of the characterising species from these communities have been recorded from the atlantic. paphia aurea, tapes decussatus, venerupis rhomboides and venerupis pullastra were represented during the atlantic, where as only venerupis pullastra is present in the limfjord today. ostrea edulis was well established during the atlantic, as seen from the species composition of the ‘køkkenmødding’ (kitchen midden) (petersen 1986a, figs 3, 4). the above-mentioned species and cerastoderma edule occurred in the infralittoral zone in large quantities most probably in the tidal zone which was the best collecting grounds for the stone age people. in a multi-lobed body of water such as the limfjord, many habitats have existed during the atlantic. however, also the development through time has been considered, as seen in the case of the marine stages in tastum sø – once the southernmost part of skive fjord (rasmussen & petersen 1980). in the northern part of the former limfjord during the atlantic, the deeper-water fauna with abra alba and corbula gibba can be demonstrated at the vust locality (petersen 1981, p. 502). the recorded atlantic communities have very much in common with the recent communities. the north sea, region 5 appendix 6 and fig. 103 twenty-seven species immigrated during the atlantic in the north sea coastal region (fig. 103), and chamelea striatula and spisula subtruncata characterise the venus community and are very common in the recent north sea region. also paphia aurea and tapes decussatus make their appearance in the north sea region during the atlantic. according to hessland (1943), tapes decussatus should immigrate to the west coast of sweden already in the boreal, while paphia aurea, venerupis rhomboides and venerupis pullastra followed in the atlantic. this is a close parallel to the recorded immigration to the limfjord region, although here following the transgression and not superjacent to older marine deposits as in the north sea. the vendsyssel region, region 6 appendix 6 and fig. 103 the dates from the vendsyssel region during the holocene are made on only a few species. however, the faunal assemblages sensu nordmann (jessen 1905) can be commented upon in the light of immigration dates observed in the neighbouring limfjord region. nordmann (jessen 1905, p. 145) operates with five assemblages (from a to e) with the following headings: a) beach deposits b) oyster banks c) deposits in coves and sounds d) deposits in fjords and sounds with muddy bottoms and no current e) lagoonal deposits a) beach deposits the first type – the beach deposits – cannot be considered in any relation to the community concept sensu c.g.j. petersen, since the dominating part of the shell material has been redeposited. however, as a geological unit, it points to a former sea level stand, albeit difficult to date, because of the allochthonous character of these deposits. among the 90 species listed from this region (southgeus bulletin no 3.pmd 09-07-2004, 09:10161 162 ern part of vendsyssel), nordmann (jessen 1905, table a) points to spisula subtruncata and fabulina fabula as being conspicuous, but other species may dominate at some localities, as seen from the table. both species pointed out by nordmann are recorded from the limfjord during the atlantic. b) oyster banks as pointed out by petersen (1918, p. 52), the so-called oyster banks in the recent limfjord have 1 or 2 specimens per m2. however, the places recorded by nordmann are located on former narrow channels where the oysters occurred in large quantities together with chlamys varia, hiatella arctica, retusa truncatula, mysella bidentata, parvicardium exiguum, paphia aurea, venerupis pullastra, bittium reticulatum, rissoa inconspicua, rissoa parva, buccinum undatum, nucula nitidosa, triphora adversa, cerastoderma edule, and hydrobia ulvae, all of which are recorded from the limfjord during the atlantic. only two species mentioned by nordmann (jessen 1905, p. 147) as being characteristic from some of the oyster banks, caecum glabrum and acmaea virginea, have their earliest record from the limfjord in the subboreal. therefore the oyster banks sensu nordmann seem to be well established in the vendsyssel region already during the atlantic, considering the dates obtained from the limfjord region. the oyster banks from the atlantic appear to be characteristic features with their high diversity of species and huge quantities of ostrea edulis not met with in present-day danish waters. this could be seen as a parallel to the fluctuation in the population of oysters observed during the last hundred years in danish waters, but should rather be connected with changes in the tidal currents which changed to a minimum during the following stage – the subboreal (petersen 1993), and put an end to the large oyster banks. c) deposits in coves and sounds from these deposits nordmann points to species such as spisula subtruncata, modiolus modiolus, thracia phaseolina, and corbula gibba as being characteristic of coves and sounds. they have all been recorded from the limfjord during the atlantic, and they represent species known from the deeper-water deposits both as epifaunal elements (modiolus modiolus) and infaunal elements as found in the c.g.j. petersen communities, the modiola and abra communities respectively. this is further demonstrated by the following species mentioned by nordmann (jessen 1905, p. 148): cerastoderma edule, parvicardium scabrum, nucula nitidosa, hiatella arctica, chamelea striatula, timoclea ovata, venerupis pullastra, fabulina fabula, tellimya ferruginosa, lunatia alderi, and retusa truncatulus. also these species have been dated back to the atlantic in the limfjord region. ostrea edulis occurs, but as stray finds among the infauna elements dominating in the above-mentioned assemblage that includes abra alba, which occurs in most of the samples, although not frequently (jessen 1905, table c). d) deposits in fjords and sounds with muddy bottoms and no current from such deposits nordmann mentioned the finds of zostera, which was a well established vegetational element in the recent limfjord, according to petersen & jensen (1911, map 1). the dominating species in this assemblage, which resembles the present-day fauna in such environments, are hydrobia ulvae, littorina littorea, littorina obtusata, rissoa membranacea, cerastoderma edule, mytilus edulis, scrobicularia plana, paphia aurea, bittium reticulatum, hinia reticulata, onoba semicostata, parvicardium exiguum, macoma balthica, and ostrea edulis, the last two species only with a few specimens. all the above-mentioned species occurred in the limfjord region during the atlantic. the littorina, rissoa, and parvicardium species might often be found on the zostera vegetation. among the dominating species also mentioned by nordmann, some have not been dated back to the atlantic (in the limfjord region) but occur in the subboreal, viz. littorina tenebrosa, akera bullata, and retusa obtusa. however, already the species recorded from the atlantic point to the so-called echinocyamus community (spärck & lieberkind 1921), although the echinoids have not been recorded by nordmann (jessen 1905). e) lagoonal deposits these deposits represent two assemblages, according to nordmann (jessen 1905, p. 150), viz. an older more open-water environment with species such as mactra stultorum, tellimya ferruginosa, chamelea striatula, fabulina fabula, ensis ensis, lunatia catena, lunatia alderi, and aporrhais pespelicani, which are mixed with faunal elements from the lagoon itself, such as hydrobia ulvae, scrobicularia plana, and mytilus edulis. in connectionwithazostera vegetation, rissoa membranacea and lacuna vincta may occur in huge quantities. geus bulletin no 3.pmd 09-07-2004, 09:10162 163 such a deposit cannot be compared to any of the petersen communities, although they play an important role in the geological setting, as was the case also with the beach deposits. in the northern and eastern part of vendsyssel, further comments will be added to the shallow-water and beach deposits with the finds of the dosinia and mya arenaria species. they have been dated to the subboreal and subatlantic respectively and are therefore commented upon later. the skagen region, region 7 appendix 6 and fig. 103 nearly all of the eight recovered species from the atlantic (fig. 103) show a deep-water fauna, which on the basis of the dominating role of the echinoids is tentatively referred to the amphiura community known from the present-day skagerrak. the final large eustatic rise took place during the late boreal – early atlantic, and the difference in isostatic rebound from 8000 b.p. between the skagen and limfjord regions is around 31 m, with the highest amount in the north (skagen). it appears that the water depth in the skagen region must have been up to 100 m during the atlantic (petersen 1981, 1991b). therefore, the occurrence of a single spisula subtruncata shell must be taken as far outside its habitat, considering that the modern depth range of this species is 0–36 m (petersen 1986c, table 2). the recorded atlantic communities and/or characteristic molluscan species are given in table 7. the subboreal stage 5000–2500 14c years b.p. the bælt sea, region 1 appendix 6 and fig. 103 there are no dated mollusc finds from the subboreal in the bælt sea region. as stated for the atlantic in this region, the bottom communities known from the present day were already established, but they included some species such as tapes and ostrea which are no longer extant in this area. however, as paphia aurea and ostrea edulis still occurred in the iron age sites – from the subatlantic – it is most probable that these species persisted there, while tapes decussatus and venerupis pullastra expired during the subboreal in the bælt sea region (petersen 1985c, fig. 5). the baltic, region 2 appendix 6 and fig. 103 there is no dated record of molluscs from the subboreal in the danish part of the baltic. therefore the change in the atlantic littorina fauna, into the lymnaea sea fauna, which occurred during the subboreal around 4000 b.p. (fredén 1980, p. 70), must be taken from observations outside denmark. the mollusc fauna from estonia shows that littorina littorea, rissoa membranacea, and scrobicularia plana persisted there until about the end of the subboreal (kessel & raukas 1979, fig. 9). the implications of this should be that these species must have been present in the danish area throughout the subboreal. in the central part of the baltic, around gotland, lymnaea peregra f. baltica re-immigrates after the 1. bælt sea mytilus m. edulis/ostrea edulis littoral macoma m. balthica/tapes spp. shallow abra a. alba deeper 2. baltic mytilus m. edulis/littorina littorea littoral macoma m. balthica shallow deeper 3. kattegat mytilus m. edulis littoral macoma m. balthica/tapes spp. shallow abra a. alba deeper 4. limfjord mytilus/modiola m. edulis/tapes spp. littoral macoma m. balthica shallow abra/venus a. alba deeper 5. north sea littoral shallow venus chamelea striatula deeper 6. vendsyssel littoral modiola m. modiolus shallow abra a. alba deeper 7. skagen littoral shallow amphiura parvicardium minimum ~ 100 m region community species depth table 7. atlantic communities and/or characteristic molluscan species geus bulletin no 3.pmd 09-07-2004, 09:10163 164 maximum of the littorina transgression (munthe 1940, p. 124). this gastropod was also present in the early, more brackish part of the littorina sea deposits in fakse bugt. in estonia (kessel & raukas 1979, fig. 9), the reappearance of lymnaea took place around 4000 b.p., implying that the salt-demanding species (littorina, rissoa, and scrobicularia) occurred together with the brackish lymnaea species throughout the later part of the subboreal! the kattegat, region 3 appendix 6 and fig. 103 the mollusc faunas recorded from the kattegat region represent only part of the total faunal complex within this large region, and have been dated only on djursland. however, this demonstrates the expiring tidal amplitude in the early part of the subboreal. the only dated immigrants to the fauna from the central part of djursland during the subboreal are onoba semicostata and littorina tenebrosa. both extended into the baltic today and tolerate brackish water. in this way they are typical for the environmental changes recorded in the marine faunas from djursland. the fauna during the atlantic was characterised by ostrea edulis, tapes decussatus, macoma balthica, and corbula gibba. bittium reticulatum was present in large quantities, but disappeared in the subboreal. also the decline in numbers of hydrobia ulvae and its replacement in equal numbers by hydrobia ventrosa speak in favour of a more brackish-water influence. the implications of the study of mollusc species on a quantitative basis in connection with 14c dates and pollen analyses confirm that the fauna during the tapes sea period was more prolific than nowadays. however, it also demonstrates as a new point of view that this applies only for the atlantic. in petersen (1993, p. 368) it is argued that the change in sedimentation rate from the atlantic to the subboreal, which has been calculated for the korup sø area on djursland, points to a lowering of the tidal range in danish waters since the atlantic. this is explained in that way that sedimentation will stay low as far as the tidal current reaches and allows halophilous species to live far up in the fjords according to the observations on the faunal record. furthermore, an older record from the mapping of the area of flaser bedding seen as a tidal bedding supports such an explanation. it was tempting to see the change from the littorina sea to the lymnaea sea in the baltic on the background of such a lowering of the tidal impact in the inner danish waters. however, as shown in the preceding section on the baltic, the change occurred around 4000 b.p. recalling the statement by c.g.j. petersen that the deposition of the tapes layers has happened in a period when the danish waters from a hydrographical point of view have been more like the north sea or the open sea than now, it is clear that a tidal impact could make the difference and explain the large oyster banks far into the roskilde fjord in north-eastern sjælland and other former fjord regions facing the kattegat region. the well-dated ertebølle coastal sites (‘køkkenmøddinger’ – kitchen middens) from all over denmark also present a large amount of ostrea edulis from the atlantic and demonstrate that the molluscan diet later in the subboreal was based on the cardium species (andersen 1991, 1995). this situation has lasted into the iron age, as seen in the shell middens from the bælt sea area (petersen 1985c, fig. 5). however, this change mostly affected the fjord complex. consequently the kattegat region still has the communities listed for the atlantic. the limfjord, region 4 appendix 6 and fig. 103 the 36 species which immigrated into the limfjord during the subboreal (appendix 6) can be considered according to their way of life, presented from the list below. age: subboreal climatic regions: asb. class gastropoda subclass prosobranchia order archaeogastropoda margarites helicinus (phipps 1774) subclass opisthobranchia order anaspidea retusa obtusa (montagu 1803) total for climatic regions asb. : 2 (5.6%) climatic regions: asbl class gastropoda geus bulletin no 3.pmd 09-07-2004, 09:10164 165 subclass prosobranchia order archaeogastropoda acmaea tessulata (müller 1776) order neotaenioglossa lacuna pallidula (da costa 1778) total for climatic regions asbl: 2 (5.6%) climatic regions: .sbl class gastropoda subclass prosobranchia order archaeogastropoda acmaea virginea (müller 1776) class bivalvia subclass pteriomorpha order pterioida delectopecten vitreus (gmelin 1791) total for climatic regions .sbl: 2 (5.6%) climatic regions: ..bl class gastropoda subclass prosobranchia order archaeogastropoda helcion pellucidum (linnaeus 1758) gibbula tumida (montagu 1803) skenea basistriata (jeffreys 1877) order neotaenioglossa littorina tenebrosa (montagu 1803) lacuna parva (montagu 1803) cingula semistriata (montagu 1808) rissoa violacea desmarest 1814 caecum glabrum (montagu 1803) aporrhais pespelicani (linnaeus 1758) order heterogastropoda epitonium turtonis (turton 1819) vitreolina philippii (rayneval & ponzi 1854) order neogastropoda hinia incrassata (ström 1768) raphitoma purpurea (montagu 1803) raphitoma linearis (montagu 1803) subclass heterobranchia order heterostropha chrysallida decussata (montagu 1803) chrysallida indistincta (montagu 1808) ebala nitidissima (montagu 1803) eulimella laevis (brown 1827) ondina divisa (j. adams 1797) ondina diaphana (jeffreys 1848) subclass opisthobranchia order anaspidea akera bullata müller 1776 class bivalvia subclass palaeotaxodonta order nuculoida nucula nucleus (linnaeus 1767) subclass pteriomorpha order pterioida palliolum striatum (müller 1776) palliolum tigerinum (müller 1776) subclass heterodonta order veneroida mactra stultorum (linnaeus 1758) lutraria lutraria (linnaeus 1758) ensis ensis (linnaeus 1758) order myoida pholas dactylus linnaeus 1758 total for climatic regions ..bl: 28 (77.8%) climatic regions: ...l class gastropoda subclass prosobranchia order neotaenioglossa alvania lactea (michaud 1830) onoba proxima (forbes & hanley 1850) total for climatic regions ...l: 2 (5.6%) total for age subboreal: 36 (14.6%) the archaeogastropoda are all (six) epifauna on seaweeds or on hard substrates. the neotaenioglossa have seven epifaunal elements mostly on seaweeds and two infauna species, of which aporrhais pespelicani is a shallow infauna animal. the heterogastropoda with epitonium turtonis and vitreolina philippii are associated with other animals, the former feeding on anemone or preying on other species and the latter being an intermittent parasite of echinoderms (fretter & graham 1982, p. 387). the heterostropha with six species are predators or external parasites. the anaspida with two species, akera bullata and retusa obtusa, are epifauna and infauna species respectively, the former on zostera in shallow water and the latter in mud or fine sand connected with the macoma community. the only nuculoida found, nucula nucleus, belongs to the shallow infauna. the pterioida with three species are referred to the epifauna, since the delectopecten vitreus is found attached with its byssus on hard substrates. the veneroida have three species which are all referred to the infauna. lutraria lutraria and ensis ensis are deep-borrowing. geus bulletin no 3.pmd 09-07-2004, 09:10165 166 the myoida with pholas dactylus bores in different substrates. when taking the above-mentioned groups of species associated with other animals, carnivores, predators, and boring species as a whole, we have three categories: the epifauna with 47.2%, the infauna with 19.4%, and other elements with 33.4% of the species immigrated during the subboreal. when the same procedure is followed for the 77 species which have been dated to the atlantic in the limfjord region, we find that the percentages for the epifauna, the infauna and other elements are 31%, 46.8% and 22% respectively, which shows that the epifauna element becomes the dominating one in the subboreal among the newcomers. this might tentatively be connected with a denser vegetation in the subboreal of sea-weed. the north sea, region 5 appendix 6 and fig. 103 nineteen species make their first appearance in the danish north sea during the subboreal. when considering their way of living and their grouping into epifauna, infauna and other elements, it appears that the groups are of equal size, i.e. five, six and eight species respectively. however, the number is too low to be used for any comparison with other regions. in the north sea region the dates of first appearance go back to the preboreal–boreal stage, showing that the initial stages were dominated by the infauna species; the preboreal–boreal: seven epifauna, 16 infauna, and four other elements; the atlantic: six epifauna, 17 infauna, and four other elements. the development of the bottom communities in the north sea region seems in this way to corroborate the changes observed in the limfjord region from the atlantic to the subboreal. these changes are in facies rather than climatic. a slightly more temperate fauna was met with during the atlantic, as mentioned earlier, and it has consequently no bearing on the observed changes. but the expiring tidal influence in the danish waters taking place in the early subboreal might have been of some importance for the environmental changes reflected through the bottom communities. the vendsyssel region, region 6 appendix 6 and fig. 103 the vendsyssel region does not give much information on the immigration of species during the holocene. however, the dosinia beds were described from this area and have been dated quite recently in the type area around strandby north of frederikshavn (nordmann 1904; petersen 1991b). the oldest date for the dosinia exoleta, which is the characteristic species for the dosinia beds, is 4240 ± 85 b.p. in 14c years (k5318). this earliest dated occurrence of dosinia exoleta corresponds to a hydrographical change in the kattegat region described by nordberg & bergsten (1988) and nordberg (1989). the demonstrated lowering of the tidal influence in the inner danish waters took place also in the early part of the subboreal. petersen (1976) pointed out that seven mollusc species hitherto known only from the dosinia beds also occur in the raised marine deposits from the western part of the limfjord, i.e. lucinoma borealis, hinia incrassata, venerupis rhomboides, abra prismatica, lutraria lutraria, pholas dactylus, and helcion pellucidum. furthermore, not only ten species from the dosinia beds are also in the deposits from the limfjord but five of them occurred already during the atlantic: gari fervensis, turritella communis, lucinoma borealis, abra prismatica, and venerupis rhomboides. these species, representing an infauna assemblage very much like the dosinia species, were also characteristic of the early holocene dominating infauna mollusc assemblage. mörner (1969, pp. 384–386, and table 1) points out that some species in the dosinia fauna occur in older layers along the swedish west coast, referring to the works by hessland (1943) and antevs (1917). however, this is not the case with the characterising dosinia species, in as much as dosinia exoleta has not been demonstrated in the studies by hessland and antevs and dosinia lincta occurs only in the younger deposits referred to the subboreal. among the 15 species listed, eight have been recorded from the limfjord, out of which epitonium turtonis, oenopota turricola, acteon tornatilis, and cylichna cylindracea have their first appearance in the atlantic, and lutraria lutraria, pholas dactylus, and alvania lactea appeared in the subboreal, whereas the occurrence of modiolus adriaticus in the limfjord has not been dated. in this way, we are left with only six species which have not been found in other regeus bulletin no 3.pmd 09-07-2004, 09:10166 167 gions outside vendsyssel older than the subboreal: pecten maximus, dosinia exoleta, dosinia lincta, gari depressa, alvania cimicoides, and trivia monacha. only the characterising species dosinia exoleta has been dated as mentioned above, and recently donax vittatus from vr. holmen in the northern part of vendsyssel, west of strandby, with the oldest date of this bivalve hitherto obtained in the danish deposits, 4240 ± 75 14c age b.p. (aar-1481). this date shows that donax characterising the highenergy shore deposits occurred in danish waters since the subboreal. the further revision of the dosinia fauna shows that only very few species are limited to the vendsyssel region both in time and space. therefore, it cannot be sustained for the danish material as mentioned by mörner (1969, p. 384) that: “the dosinia layers contains a great number of new boreo-lusitanic immigrants” of the molluscs entirely belonging to the dosinia layers according to nordmann (1904), only three species, out of the 26 species mentioned, do not occur in danish waters today, according to jensen & knudsen (1995), viz. trivia monacha, gari depressa, and alvania cimicoides. only trivia monacha seems to be purely lusitanian, since a recent distribution to the north sea is questioned by fretter & graham (1981, p. 329), and there is no record from scandinavia. the other two species have a boreo-lusitanian distribution. the skagen region, region 7 appendix 6 and fig. 103 the 23 species, out of which only vitreolina collensi is purely lusitanian, can be evaluated according to their way of life. vitreolina collensi belongs together with aclis minor and melanella alba to the eulimacea, which are probably associated with echinoderms (fretter & graham 1982, p. 397). the eulimacea, together with the epitoniacea, usually prey on anthozoans. the heterostropha, including the family pyramidellidae which lives ectoparasitically on other marine organisms (jensen & knudsen 1995), are here represented by eulimella scillae. finally within the veneroidae, mysella bidentata and tellimya ferruginosa are commensals on echinoids, but can also be found free-living. for the rest of the 23 species found during the subboreal in the skagen well, it applies that 16 species belong to the infauna, including onoba vitrea which tends “to live in muddy places, often so muddy that one wonders how the animals keep the mantle cavity clear” (fretter & graham 1978b, p. 170). it appears from the above-mentioned dates based on type of bottom-dwelling animals that the fauna belongs to the deep-level sea bottoms which goes well together with the turritella-venus communities. the recorded subboreal communities and/or characteristic molluscan species are given in table 8. the subatlantic stage 2500– 14c years b.p. the configuration of land and sea in the danish realm was very close to that of today. the isostatic movements during this time span have been so small that they did not affect the general outline (petersen 1991b). however, the coastal development, in the formation of simplified coastline and spits especially in the west towards the north sea and in the north facing the skagerrak and the kattegat respectively, still affected the contour of the land. 1. bælt sea mytilus m. edulis littoral macoma m. balthica/ostrea/paphia shallow abra a. alba deeper 2. baltic mytilus m. edulis littoral macoma m. balthica shallow deeper 3. kattegat mytilus m. edulis littoral macoma m. balthica shallow abra a. alba deeper 4. limfjord mytilus/modiola m. edulis littoral macoma m. balthica shallow venus/abra a. alba deeper 5. north sea littoral shallow venus chamelea striatula deeper 6. vendsyssel donax d. vittatus littoral dosinia d. exoleta shallow deeper 7. skagen littoral shallow venus/turritella t. communis deeper table 8. subboreal communities and/or characteristic molluscan species region community species depth geus bulletin no 3.pmd 09-07-2004, 09:10167 168 for the main part of the danish waters, the recent marine bottom communities were established, some of them already since the atlantic, although the few characterising tapes and ostrea species are no longer extant in great numbers or have totally vanished from the danish seas. therefore, the actual map of the petersen (1914, 1918) bottom communities as seen today will be commented upon in relation to the few, but important changes observed during the subatlantic, region by region. the bælt sea, region 1 appendix 6 and fig. 103 the bottom communities mapped from the bælt sea region comprise the macoma balthica community in the shallow-water area and the abra alba community in deeper water (petersen 1918), the latter community with tridonta borealis and macoma calcarea. the former has been recorded from the subfossil finds but not dated, while the latter has a dated occurrence back in the atlantic and is considered part of the abra alba community as a deep infauna element. considering the present distribution of the astarte species, it is most probable that tridonta borealis invaded the bælt sea and the baltic already in the early holocene along with the transgression in the early atlantic. the gregarious occurrences of ostrea edulis recorded from the atlantic in the bælt sea region vanished in the iron age (petersen 1985c). this species is no longer found in the bælt sea region, nor is paphia aurea, which also occurred at the iron age sites too (petersen 1985c, fig. 5). the steady occurrence of ostrea edulis since the atlantic, although in reduced numbers, might have led to an experiment in cultivating oysters south of lolland in the fehmern bælt (winther 1876, p. 114), although an unsuccessful one. the distribution of oysters within the danish waters seems to have changed very much right up to the present day, with many records from the 19th century of oyster banks from places where no records are found today (kröyer 1837; seaman & ruth 1997). the baltic, region 2 appendix 6 and fig. 103 the macoma balthica community covers the whole area of the baltic, implying that in this area the otherwise shallow-water bivalve extends into greater depths – more than 50 m (petersen 1918). also in this area, tridonta borealis and tridonta elliptica have been found in great quantities east of bornholm but not dated. however, according to johansen (1916, fig. 5), tridonta borealis and tridonta elliptica are recorded only from areas with a salinity of more than 10‰, but zenkevitch (1963, p. 338, fig. 167) points to many finds further to the north in the baltic, where the salinity is lower. as stated earlier, the change from the littorina sea stage to the lymnaea sea stage took place during the subboreal. the present situation with a mya sea stage – a term established by munthe (1894) – took place at a very late date. munthe (1894, p. 14) said: “since mya arenaria is an easily identified and characteristic species in the present baltic it seems suitable to call the present time the ‘mya-time’ or ‘mya arenaria-time’ in opposition to ‘littorina-time’ etc.” the kattegat, region 3 appendix 6 and fig. 103 among the subfossil species both, dated and undated, no records of chlamys striatula and turritella communis are found. these characterising species for the venus and deeper venus communities respectively have a wide extension on the map by petersen (1918) in the kattegat region. also the deeper-water epifaunal elements – characterised by modiolus modiolus – are missing in our subfossil record. only the macoma balthica and abra alba communities are recognised in the subfossil material. however, the development in the skagen well sequence to the north in the kattegat region of mollusc species reveals the venus-turritella communities and can be taken as part of the development in the central kattegat region not sampled at the time of this study on molluscs. in the present day northern kattegat, stray specimens of ostrea edulis have been recorded (jensen & knudsen 1995, p. 40). otherwise among the more spectacular tapes sea species, tapes decussatus, dated from the atlantic, and paphia aurea, not dated but occurgeus bulletin no 3.pmd 09-07-2004, 09:10168 169 ring in the subfossil fauna, have disappeared from the danish waters. the limfjord, region 4 appendix 6 and fig. 103 this region has been studied in more detail, regarding the recent fauna, than the other regions, and references can be made also to jensen (1919). the abra community is here divided into three associations, i.e. nucula-corbula, abra-solen, and abrasolen-mya associations. in more shallow water the abra community is replaced by the macoma balthica community. all of these communities are recorded by their mollusc species in the subfossil fauna, here including the mya arenaria in the subfossil assemblage. however, the only dated subfossil immigrating species from thesubatlantic isdonaxvittatus,whichappearedaround 2000 years b.p. in the northernmost part of the limfjord region in the beach ridges, around 1000 years before the closing of the western and northern entrance to the limfjord. the closing of the entrances from the north sea and skagerrak changed the limfjord region into a freshwater basin between a.d. 1200 and 1825, however, with periods of saltwater influence (kristensen et al. 1995). a comparison between the subfossil fauna before the freshwater stage and the recent one after the north sea broke through in the western part of the limfjord in 1825 shows that the subfossil fauna had only a slightly higher affinity to more temperate water than the recent one. also in this place the tapes, venerupis and paphia species make the difference, in the way that only venerupis pullastra finds its way back to the region after 1825. paphia aurea, however, has a dated occurrence from the same deposits as donax vittatus to 1910 ± 100 14c years b.p. (petersen 1976). it is seen that paphia aurea in this region, as in the bælt sea, has a record up into the iron age before it became extinct in the danish waters. ostrea edulis repopulated the limfjord region after 1825 and reached a wide extension in this region already in the second half of the 19th century (collin 1871). however, the population has suffered from strong fluctuations not only in the limfjord but also in other danish waters, as shown on the map by kröyer (1837). spärck in several papers on the biology of oysters (ostrea edulis), published in reports of the danish biological station, also discussed the fluctuations in the nw european population of oysters (spärck 1950, pp. 43–45). spärck reached the conclusion that the summer temperature of the water was crucial, both being too low and too high, which affected the oyster in its reproduction and in food supply respectively. furthermore, severe winters might affect the population, although less than the summer temperatures. however, these changes did not mean a total disappearance of the oyster, but only a reduction to such a level that the industrial exploitation had to stop. when taking into account the many studies on the population of oysters, one could use the results in a general conclusion on the variations found in the whole population of molluscan species, especially for the group having their northern limit within the boreal region: even small variations in the climate may influence the size of the population. also the environmental changes as shown within the danish area during the holocene, such as the lowering of the tidal amplitude in the early part of the subboreal, had a severe influence on the populations in the inner danish waters. here again, oysters can be taken as an example by the termination of the huge oyster banks known from the atlantic. spärck (1950, p. 44) draws attention to the oyster banks in holland and the british isles, where the density of the population is far greater because of the tidal movements. however, not only the hydrographical changes through time, in the tide, but also the coastal evolution, such as the formation of simplified coastline and spits, play an important role in the distribution and new finds of molluscan species. the north sea, region 5 appendix 6 and fig. 103 only eight species have been recorded as immigrants during the subatlantic. however, two of them, donax vittatus and dosinia lincta, deserve special attention. only parvicardium ovale and dosinia lincta have their first dated appearance. the other species, except donax vittatus with occurrence in the subboreal, have been recorded from the atlantic at various places listed in appendix 6. in referring to the c.g.j. petersen bottom community map covering also the north sea, the macoma geus bulletin no 3.pmd 09-07-2004, 09:10169 170 balthica and the venus communities are found in the danish north sea coastal region, the former in bays and off the southern part of the west coast (petersen 1914), the latter around the westernmost part of the limfjord and the jydske rev wnw of the bovbjerg coastal cliff (petersen 1994a). a landscape like the limfjord of today was found 75 km towards wnw in the area of the jydske rev. following the transgression in the early part of the holocene, the glacial landscape in an area of the present jydske rev was eroded and the high-energy coast approached the appearance of the present one. in the northern part forming an erosion coast and in the southern part at blåvands huk an aggradation coast, both characterised by the presence of donax vittatus. in the southern part the aggradation started around 800 bc some 2000 m east of the present coastline (petersen 1994a, p. 24) as seen from the dating of donax vittatus to 2620 ± 75 b.p. 14c years (aar-1480) off the inland cliff at grærup (fig. 1). at bovbjerg, the strata with donax vittatus in the agger spit are dated to 410 ± 65 b.p. (petersen 1985a). the formation of the spits closing the former bays on the jylland west coast is a consequence of the formation of a simplified coast. further to the north, donax vittatus from kovad bro in the northernmost part of the limfjord, 6 km inland, gave a date of 1910 ± 100 b.p. (petersen 1976), showing that the beach progressed 6 km during approximately 2000 years (petersen & andreasen 1989, fig. 1). it is tempting to introduce the idea that the enormous change in the land–sea configuration in the eastern part of the north sea affected the tidal currents in the inner danish waters. this could possibly have occurred when most of the jydske rev formation was eroded to such a level that the tidal current from the south was no longer braked and consequently the present-day interference with the tidal current came into existence. it is the interference between the two tidal currents in the skagerrak today that makes the tidal amplitude small in the inner danish waters (nielsen 1939; kuenen 1950). dosinia lincta has been dated (870 ± 110 b.p. 14c years) in the jydske revsand formation in the vibrocore 562001 around 75 km off the coast of jylland at a depth of 32 m (petersen 1994a, p. 18, fig. 3). the assemblage from these strata comprises spisula subtruncata, phaxas pellucidus, fabulina fabula, chamelea striatula, dosinia exoleta, corbula gibba, cochlodesma praetenue and thracia phaseolina, most of them characterising the jydske revsand formation. the vendsyssel region, region 6 appendix 6 and fig. 103 the coastal development in the eastern part of the region facing the kattegat takes place in the form of migrating bars (schou 1949, fig. 17b), the so-called rimmer and doppe system sensu jessen (1905). the venus and the macoma balthica communities are found in the coastal zone, the former dominating in the northern part, whereas the latter forms a small area between shore and the venus community to the south towards the entrance to the limfjord at hals (petersen 1918). in the north at strandby, locus typicus of the dosinia beds, the layers with dosinia exoleta are superposed by a layer characterised by spisula subtruncata. these beds with spisula in great quantities were dated to 2640 ± 75 (petersen 1991b), the end of the subboreal, and at a level of 4.2 m a.s.l. this corresponds to a stage in the development of the skagen spit up to 4 km south of højen, where the beach ridges have an elevation of 5 m a.s.l. in the southern part of the vendsyssel region around hals another of the faunal elements of the dosinia beds – lutraria lutraria – has for long been regarded as extinct (petersen 1992). however, “from 1990 onwards live specimens have been collected regularly near frederikshavn and on the skagerrak-coast” (jensen & knudsen 1995, p. 43). also many shells of lutraria lutraria were found along the shore south of jerup halfway between frederikshavn and aalbæk. the immigration of mya arenaria cannot be taken as an indication of changes in climate, as this species mainly belongs to the boreal region and has been transferred by man from north america. what made the find so important has a more historical than geological bearing, namely that the dates obtained from the sampling at jerup demonstrated that the american softshell clam (mya arenaria) predated columbus’ voyage in 1492, having an age of a.d. 1245–1295 at ± 1 s.d. this led to the conclusion that the vikings were better candidates than columbus to be the first to find north america (petersen et al. 1992b). the significance of changes in facies is clearly demonstrated in the next and final section describing the subatlantic faunal development in the skagen region. geus bulletin no 3.pmd 09-07-2004, 09:10170 171 the skagen region, region 7 appendix 6 and fig. 103 the subatlantic molluscan fauna from the skagen well comprises 68 species with 75% belonging to the boreolusitanian region and only one lusitanian species, vitreolina collensi. however, the more interesting fact from the younger part of the skagen sequence is the total lack of macoma balthica. in this way it presents the finest resemblance with the recent bottom community map (petersen 1918), and shows that the macoma balthica community disappears in the northern part of the east coast of jylland. this means that during the last stage of the spit formation at the site of the present skagen animals, from the venus community dominated along shore. this is alsodocumentedby thehighamount of infauna elements, with 35 out of the 68 species recorded. furthermore, some of the 11 epifauna gregarious species usually connected with the vegetation can be excluded, since they occur only as stray finds, viz. lacuna pallidula, rissoa violacea, and bittium reticulatum, as discussed earlier. the rest of the molluscs (22 species) are carnivores, predators, external parasites, and commensals. turning these figures into percentages, the epifauna species amount to 16.2% and the infauna to 51.5%. comparing this with the limfjord region where an equal number of species have been found during the atlantic and the subboreal, it appears that the number of infauna species from the atlantic to the subboreal falls from 46.8% to 19.4% and the epifauna elements rise from 31.2% to 47.2% in the limfjord. counting the limfjord region as an inner danish water today, it is worth noticing that during the atlantic the situation was much more like the ‘open’ waters as seen in the skagen figures. considering that the tidal amplitude really was lowered in the early part of the subboreal, this would to some extent explain the observed changes in the limfjord from the atlantic to the subboreal. the recorded subatlantic communities according to the maps by petersen (1914,1918) with characteristic molluscan species are shown in table 9. table 9. subatlantic communities according to the bottom community maps* with characteristic molluscan species 1. bælt sea mytilus m. edulis littoral macoma m. balthica/÷ostrea/paphia shallow abra a. alba deeper 2. baltic mytilus m. edulis littoral macoma m. balthica/+mya arenaria shallow macoma m. balthica deeper 3. kattegat mytilus m. edulis littoral macoma m. balthica shallow abra/venus a. alba deeper 4. limfjord mytilus/modiola m. edulis littoral macoma m. balthica/÷paphia aurea shallow abra a. alba deeper 5. north sea donax d. vittatus littoral macoma m. balthica shallow venus chamelea striatula deeper 6. vendsyssel littoral macoma m. balthica shallow venus chamelea striatula deeper 7. skagen donax d. vittatus littoral spisula s. subtruncata shallow venus chamelea striatula deeper * petersen (1914, 1918). region community species depth geus bulletin no 3.pmd 09-07-2004, 09:10171 geological survey of denmark and greenland bulletin 13, 2007, 09-12 9 the upper cretaceous – danian chalk may be considered to be the economically most important rock type in denmark. onshore it constitutes an important groundwater aquifer and it is also quarried for e.g. building materials and paper production. offshore the chalk reservoirs contain more than 80% of the oil and gas produced in denmark (fig. 1). during the last few years efforts have therefore been made to map this important succession in the danish and adjoining areas (vejbæk et al. 2003). the stratigraphic interval mapped comprises the chalk group of cenomanian to danian ages and its stratigraphically equivalent units (fig. 2). the north-eastern limit of the chalk group is determined by neogene erosion. the limits of the map to the west and south were mainly determined by the amount of available data. data base the comprehensive data base comprises high-resolution and conventional 2-d and 3-d reflection seismic data as well as published maps (e.g. britze et al. 1995; hommel 1996; ottesen et al. 1997; jensen 1998; kramarskiej 1999; balds chuhn et al. 2001; stoker 2005). more than 500 deep wells and numerous onshore water wells have provided control for the mapping. this is especially relevant for the mapping where the top chalk is immediately overlain by the neogene (fig. 3). in these areas in particular, mapping was based on high-resolution seismic data. depth conversion depth conversion was undertaken by using depth-dependent velocity functions, where the velocity v at depth z is given by: v=v0 + dv + k×z where v0 is the surface velocity, dv is a variation of the surface velocity and k is the gradient of velocity increase with depth (table 1; e.g. japsen 1998, 1999). the surface velocity variation is typically mapped on the basis of well data and may reflect lateral facies changes, burial anomalies or excess fluid pressures. chalk depth structure maps, central to eastern north sea, denmark ole v. vejbæk, torben bidstrup, peter britze, mikael erlström, erik s. rasmussen and ulf sivhed © geus, 2007. geological survey of denmark and greenland bulletin 13, 9–12. available at: www.geus.dk/publications/bull fig. 1. hydrocarbon accumulations in the north sea with chalk fields highlighted. 10 the cenozoic velocity model consists of a single layer on shore denmark and two layers offshore. the division be tween the two layers is taken at the ‘near top middle miocene marker’ that corresponds approximately to the top of the over-pressured section (upper and lower post chalk group in table 1). the parameters for these layers were taken from britze et al. (1995) and japsen (1999, 2000) who derived a similar but segmented model for the chalk group. since the parameters are based on a large well data base from the entire north sea (e.g. japsen 2000), they are applicable to most of the north sea. notes about the maps in some areas where the neogene lies directly on the top chalk seismic horizon, the erosional truncation of the chalk group is negligible. this occurs around copenhagen, in northern sjæl land and in south-western scania, where minor outliers of selandian deposits document the former extent of the chalk group. the occurrence of palaeo gene sediments offshore poland also indicates that erosion of the chalk group is generally not very deep in the western baltic outside the main inversion zones (fig. 3). in norwegian waters, however, extensive neogene erosion has occurred. the erosion in these areas is sufficiently deep for lower cretaceous deposits to subcrop the base of the neogene. outside these areas the chalk group generally has a larger areal extent than the lower cretaceous. (fig. 3). a general increase in thickness of the chalk group is found west of the sorgenfrei–tornquist inversion zone. a north-eastward increase in thickness is also found in the areas unaffected by neogene erosion offshore southern norway, sug gesting the presence of similar depocentres on the flanks of inversion zones. thus, inversion may also have occurred in the south-western coastal areas of norway. fig. 2. lithostratigraphic correlation for the upper cretaceous – danian succession as map ped in this paper. based on deegan & scull (1977), isaksen & tonstad (1989), johnson & lott (1993) and schiøler et al. (2007) with additions modified from surlyk et al. (2003) and sivhed et al. (1999). ha. mb, hansa member; kbh. mb, københavn member; kh. mb, kyrkheddinge member; lk. mb, landskrona member. facing page: fig. 3. simplified structure maps of the chalk group and equivalent deposits. a, depth to top chalk group; b, depth to base chalk group and c, isopach. grey shadings in a and c indicate where the lower cretaceous subcrops quaternary sediments (i.e. where the chalk group has been totally removed by erosion). cph, copenhagen. pdf versions of the maps with more detail are available from www.geus.dk/publications/bull/nr13/index-uk.htm 11 12 authors’ addresses o.v.v., t.b., p.b. & e.s.r., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ov@geus.dk m.e. & u.s., sveriges geologiska undersökning, kiliansgatan 10, s-223 50, lund, sweden. hydrocarbon aspects the chalk group in the central graben area is an important reservoir and migration path for oil and gas. it is the most important oil-producing interval in denmark and is also a major contributor to oil and gas production in norway and the netherlands, while production from the chalk group is still insignificant in the uk sector (fig. 1). traps within the chalk group range from inversion-generated anticlines (e.g. the valhall, roar, tyra and south arne fields), over salt domes with some degree of inversion overprint (e.g. the dan, eko fisk and svend fields) to salt diapirs (e.g. the skjold and harald fields). stratigraphic traps may also play a major role (e.g. the halfdan and adda fields). these traps owe their existence to a combination of over-pressuring and early hydrocarbon invasion to preserve the quality of their reservoirs de spite the great depths to which they have been buried (e.g. anderson 1999; vejbæk in press). their position directly above the main upper jurassic source rock also seems to be a necessary condition for their existence (e.g. anderson 1999; sur lyk et al. 2003), since the generally very low permeability of the chalk precludes long-distance migration and even keeps accumulations in hydrodynamic dis-equilibrium (e.g. dennis et al. 2005; vejbæk et al. 2005). references anderson, j.k. 1999: the capabilities and challenges of the seismic me thod in chalk exploration. in: fleet, a.j. & boldy, s.a.r. (eds): pe tro leum geology of northwest europe. proceedings of the 5th con ference, 939–947. london: geological society. baldschuhn, r., binot, f., fleig, s. & kockel, f. 2001: geotektonischer atlas von nordwest-deutschland und dem deutschen nordsee-sektor. geologisches jahrbuch reihe a 153, 88 pp. + 3 cd-roms. britze, p., japsen, p. & andersen, c. 1995: the danish central graben: top chalk and the post chalk group. two-way travel time and depth and interval velocity, 1:200 000. geological survey of denmark map series 47, 5 pp. + 3 maps. deegan, c.e. & scull, b.j. 1977: a proposed standard lithostratigraphic nomenclature for the central and northern north sea. report of the institute of geological sciences 77/25, 35 pp. (also published as norwegian petroleum directorate bulletin 1). dennis, h., bergmo, p. & holt, t. 2005: tilted oil-water contacts: modelling effects of aquifer heterogeneity. in: doré, a.g. & vining, b.a. (eds): petroleum geology: north-west europe and global perspectives. proceedings of the 6th petroleum geology conference, 145–158. london: geological society. hommel, v. 1996: structural evolution of the rønne–kolobrzeg graben area. unpublished m.sc. thesis, 82 pp. university of copenhagen, denmark. isaksen, d. & tonstad, v. 1989: a revised cretaceous and tertiary litho stratigraphic nomenclature for the norwegian north sea. norwegian petroleum directorate bulletin 5, 59 pp. japsen, p. 1998: regional velocity-depth anomalies, north sea chalk; a record of overpressure and neogene uplift and erosion. american associa tion of petroleum geologists bulletin 82, 2031–2074. japsen, p. 1999: overpressured cenozoic shale mapped from velocity anomalies relative to a baseline for marine shale, north sea. petro leum geoscience 5, 321–336. japsen, p. 2000: fra kridthav til vesterhav, nordsøbassinets udvikling vurderet ud fra seismiske hastigheder. geologisk tidskrift 2002/2, 36 pp. jensen, s.k. 1998: structural development of the sorgenfrei–tornquist zone and adjacent areas, norwegian north sea sector. unpublished m.sc. thesis, 88 pp. university of aarhus, denmark. johnson, h. & lott, g.k. 1993: cretaceous of the central and northern north sea. in: knox, r.w. o’b. & cordey, w.g. (eds): lithostrati graphic nomenclature of the uk north sea, 2, 169 pp. nottingham: british geological survey. kramarskiej, r. (ed.) 1999: geological map of the baltic sea bottom without quaternary deposits, 1: 500 000. gdansk-warszawa: panstwowy instytut geologiczny. ottesen, d., bøe, r., longva, o., olsen, h. a., rise, l., skilbrei, j. r. & thorsnes, t. 1997: geologisk atlas – skagerrak. atlas over kvartære avleiringer, bunnsedimenter, berggrunn og batymetri i norsk sektor av skagerrak. norges geologiske undersøkelse rapport 96.138, 55 pp. schiøler, p. et al. 2007: lithostratigraphy of the palaeogene – lower neogene succession of the danish north sea. geological survey of denmark and greenland bulletin 12, 77 pp. sivhed, u., wikman, h. & erlström, m. 1999: beskrivning till berggrunds kartorna 1c trelleborg nv och so samt 2c malmö sv, so, nv och no. (sgu serie af 191, 192, 193, 194, 196, 198, skala 1:50 000). uppsala: sveriges geologiska undersökning, 143 pp. stoker, s. 2005: chalk play of the uk central graben, 11 pp. department of trade and industry workshop, aberdeen, 23 november, 2005. poster (available on: http://www.og-mrp.com/dissemination/work shops/ukcs/posters/ukcs_chalk_play_bgs.pdf). surlyk, f., dons, t., clausen, c.k. & higham, j. 2003: upper cretaceous. in: evans, d. et al. (eds.): the millennium atlas: petroleum geology of the central and northern north sea, 213–233. london: geological society. vejbæk, o. in press: on dis-equilibirium compaction as the cause for the cretaceous–paleogene over-pressures in the danish north sea. ame r ican association of petroleum geologists bulletin. vejbæk, o.v., bidstrup, t., britze, p., erlström, m., rasmussen, e.s. & sivhed, u. 2003: chalk structure maps of the central and eastern north sea. danmarks og grønlands geologiske undersøgelse rapport 2003/106, 55 pp. vejbæk, o.v., frykman, p., bech, n. & nielsen, c.m. 2005: the history of hydrocarbon filling of danish chalk field. in: doré, a.g. & vining, b.a. (eds): petroleum geology: north-west europe and global perspectives. proceedings of the 6th petroleum geology conference, 1331–1346. london: geological society. geological survey of denmark and greenland bulletin 6, 1-4 1 geological survey of denmark and greenland bulletin 6 · 2004 east greenland caledonides: stratigraphy, structure and geochronology edited by a.k. higgins and feiko kalsbeek geological survey of denmark and greenland ministry of the environment geus bulletin 6.pmd 10-02-2005, 09:531 2 geological survey of denmark and greenland bulletin 6 keywords caledonides, east greenland, geochronology, stratigraphy, structure. cover west-dipping white and rusty brown quartzites of the lower cambrian slottet formation, at slottet in the eleonore sø foreland window, resting unconformably on dark clastic sediments of the palaeoproterozoic eleonore sø complex. the summit of slottet (1933 m high) is 600 m above the glacier surface. chief editor of this series: adam a. garde scientific editors of this volume: a.k. higgins and feiko kalsbeek editorial secretaries: esben w. glendal and birgit eriksen critical readers: arild andresen (norway), brian chadwick (uk), lars clemmensen (denmark), clark r.l. friend (uk), david g. gee (sweden), john s. peel (sweden), john e. repetski (usa), minik rosing (denmark), rob strachan (uk) and martin whitehouse (sweden) illustrations: helle zetterwall digital photographic work: benny m. schark and jakob lautrup graphic production: knud gr@phic consult, odense, denmark printers: schultz grafisk, albertslund, denmark manuscripts submitted: 10 november 2003 final versions approved: 9 june – 14 september 2004 printed: 30 december 2004 isbn 87-7871-148-7 geological survey of denmark and greenland bulletin the series geological survey of denmark and greenland bulletin replaces geology of denmark survey bulletin and geology of greenland survey bulletin. citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 6, 93 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk or geografforlaget aps rugårdsvej 55, dk-5000 odense c, denmark phone: +45 63 44 16 83, fax: +45 63 44 16 97, e-mail: go@geografforlaget.dk © danmarks og grønlands geologiske undersøgelse (geus), 2004 geus bulletin 6.pmd 10-02-2005, 09:532 3 contents preface a.k. higgins . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4 lower palaeozoic stratigraphy of the east greenland caledonides m.p. smith, j.a. rasmussen, s. robertson, a.k. higgins and a.g. leslie . . . . . . . . . . . . . . . . . . . . . . . . . 5 the neoproterozoic rivieradal group of kronprins christian land, eastern north greenland m.p. smith, a.k. higgins, n.j. soper and m. sønderholm . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 29 the caledonian thin-skinned thrust belt of kronprins christian land, eastern north greenland a.k. higgins, n.j. soper, m.p. smith and j.a. rasmussen . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 41 palaeoproterozoic age of a basement gneiss complex in the charcot land tectonic window, east greenland caledonides k. thrane . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57 reconnaissance pb-pb dating of single mineral phases by the step-leaching method: results from the caledonides of east greenland k. thrane . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 67 the eleonore sø and målebjerg foreland windows, east greenland caledonides, and the demise of the ‘stockwerke’ concept a.k. higgins and a.g. leslie . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 77 geus bulletin 6.pmd 10-02-2005, 09:533 4 preface the east greenland caledonides extend from 70° to 81°30′n, and have been the subject of a series of regional mapping programmes between 1968 and 1998. the entire orogen is now covered by five published 1:500 000 geological map sheets. the six papers in this bulletin concern a variety of topics relating mainly to kronprins christian land (79°–81°30′n) and the kong oscar fjord region (72°–75°n). the paper by smith et al. on lower palaeozoic stratigraphy proposes amendments to several stratigraphical units that occur in kronprins christian land and nearby lambert land. in the kong oscar fjord region, two new formations are defined for quartzite and limestone/dolostone units that crop out in foreland windows, and the lower palaeozoic succession of the fjord region of east greenland is formally placed in the kong oscar fjord group. the second paper by smith et al. describes and formally defines the neoproterozoic rivieradal group of kronprins christian land. the paper by higgins et al. analyses the thinskinned fold-and-thrust belt that marks the transition between foreland and orogen in kronprins christian land, and presents a balanced cross-section restoration. the two geochronological papers by thrane report the results of ion microprobe zircon analyses from orthogneisses in the charcot land window (72°n), and results of reconnaissance pb-pb dating by the stepleaching method. the final paper by higgins & leslie reviews the history of geological research in the eleonore sø and målebjerg areas of the kong oscar fjord region (72°– 75°n). recognition that the two areas are part of the caledonian foreland implies that the two thrust sheets structurally overlying the eleonore sø and målebjerg windows have large displacements (~ 100 km each), and that the ‘stockwerke’ concept of the orogen that focused on in situ vertical movements can finally be laid to rest. a.k. higgins geus bulletin 6.pmd 10-02-2005, 09:534 geological survey of denmark and greenland bulletin 7, 2004, p 17-20 17 seismic inversion has become a standard tool for porosity prediction in chalk exploration and field development. avo (amplitude versus offset) processing for fluid prediction is, however, still not widely applied in chalk. forward modelling may help to assess the value of acquiring these data, as well as support subsequent interpretation. this paper presents a forward modelling tool to simulate seismic response resulting from changes in degree of compaction and fluid contents in profiles. modelling is carried out with software developed by norsk hydro a/s (the compound model builder), where the geological survey of denmark and greenland has supplied special capabilities to model north sea chalk. the forward modelling is illustrated with a seismic line straddling the dan and halfdan chalk fields (fig. 1). hydrocarbons are found in the uppermost danian and maastrichtian chalk where porosity may exceed 30%. production takes place from wells centred on an anticline in the dan field, and although structural closure is lacking for the halfdan field, a surprising laterally continuous oil column is found between the two fields (e.g. jacobsen et al. 1999; albrechtsen et al. 2001). the modelling target is the effect on seismic data of such oil occurrences. the rather subtle impact on rock properties requires very realistic and detailed modelling. fortunately, the pelagic deep-water origin of the chalk makes it very uniform laterally, such that compaction effects and fluid changes handled by the modelling tool account for lateral seismic changes, whereas well data support vertical changes. forward modelling approach prediction of compaction effects starts with modelling porosity change, and subsequently the seismic properties of bulk density and pand s-wave velocities are calculated. these predictions account for changes in acoustic properties away from wells, as in synclines and anticlines. empirical compaction laws are applied where excess fluid pressure is accounted for in a simplistic way (vejbæk et al. in press b). an exponential decay of porosity (φ) with depth (z) is assumed: geological survey of denmark and greenland bulletin 7, 17–20 (2005) © geus, 2005 forward modelling of seismic response from north sea chalk ole valdemar vejbæk and rasmus rasmussen 5°05′e m-9x 1900 1950 1950 1900 2000 2075 2075 2050 2050 1900 2050 55°30′n nana-1xp 4°30′e 5°00′e 5600′n 55°45′n 55°30′n 5615′n below 1600 no data 1600–1700 1700–1800 1800–1900 1900–2000 2000–2100 2100–2200 2200–2300 2300–2400 2400–2500 2500–2600 2600–2700 2700–2800 2800–2900 2900–3000 3000–3100 3100–3200 3200–3300 3300–3400 above 3400 metres 25 km 2 km hc column 20 m0 4°00′e3°30′e halfdan field dan field north sea denmark a b fig. 1. a: map showing top chalk depth structure for the danish north sea area and location of dan and halfdan fields. producing oil and gas fields are shown in green and red, respectively. b: location of the simulated seismic line from the halfdan to the dan field (see fig. 3). the zig-zag geometry follows trace and line directions in the 3d seismic survey. colour bar shows net height of danian and maastrichtian hydrocarbon column. top chalk depth contours in 25 m interval. φ = φ0 . e -az where each layer has a surface porosity (φ0) and a is the decay parameter (sclater & christie 1980; jensen et al. 1985). further developments to account for overpressure and more complex chalk porosity decay are given by vejbæk et al. (in press b). deviation of average porosity from the normal compaction trend is due to overpressure and/or early hydrocarbon invasion which stops porosity decay (fig. 2a). the method yields good predictions of excess fluid pressures in the absence of hydrocarbon preserving effects. these predictions rely on good calibration of the compaction parameters, and on the presumably very late onset and modest subsequent dissipation of overpressure in the studied region. the chalk compaction trend is based on the revised normal velocity–depth trend of japsen (1998) converted into a porosity–depth trend (vejbæk et al. in press b). the decompaction parameterisation addresses only thick chalk successions and does not account for small-scale variations in porosity as on well logs. adjusting the surface porosity for each log sample after overpressure correction retains these details. each sample thus compacts or decompacts along a slightly different porosity decay path. local-scale variations are thus assumed to be inherited from time of deposition. this causes porosity variations to increase during decompaction and decrease during compaction. this behaviour complies with the general observation that highly reflective chalk often correlates with high porosity (britze et al. 2000). before well data are entered into the model builder, they are normalised to an arbitrarily chosen depth (2100 m in the case shown) using the approach described. following interpolation along the interpreted horizons in the model builder, porosity is then restored to the appropriate depth along the profile. the process honours observed porosity profiles at well sites and accounts for depth changes along the profile. following compaction/decompaction, sonic and density logs are corrected. the corrections are based on a modified hashin-shtrikman model as proposed by walls et al. (1998) for ekofisk field data (see vejbæk et al. in press b). the model describes how bulk and shear moduli change with porosity in an interval between zero porosity and a chosen maximum porosity (fig. 2b, c). stiffer chalk types plot closer to the modified upper hashin-shtrikman (muhs) model, and softer chalk types plot closer to the lower bounding limit. during compaction, changes in bulk moduli are adjusted according to the modified hashin-shtrikman model. changes in compaction cause changes in hydrocarbon saturation as given by the saturation model. calculation of the fluid effects involves modelling of the saturation distribution and resulting effects on seismic properties of the rock. hydrocarbon saturation is modelled with the equivalent radius method (eqr; engstrøm 1995). this model yields drainage equilibrium saturation based on estimates of capillary pressure in relation to capillary entry pressure. the capillary entry pressures in the chalk are considerable, and distances between free water levels (where oil and water pressure are equal) and the oil–water contacts are usually measured in tens of metres. hydrocarbon saturation functions for danian and maastrichtian chalk differ slightly, with danian chalk being less favourable. changes in acoustic properties caused by changes in fluid content are calculated by fluid substitution (gassmann 1951). the gassmann theory has been shown to apply to laboratory elastic measurements on chalk and therefore also applies to log and seismic data (fabricius et al. 2002; japsen et al. 2002, 2005). all the compaction changes and the seismic property calculation including fluid effects are combined into one routine in the compound model builder. in order to investigate various scenarios of hydrocarbon distribution, editing of the free water level (fwl) is sufficient. 18 0 10 20 30 0 0.2 0.4 0 20 40 60 80 0 0.2 0.4 porosityporosity porosity d ry s h ea r m o d u lu s (g p a) d ry b u lk m o d u lu s (g p a) 0 0.2 0.4 0.6 0.8 2000 1000 0 a b c d ep th ( m ) this paper sclater & christie (1980) subsidence path for overpressured chalk a b fig. 2. a: schematic compaction behaviour of chalk (black curve) with the sclater and christie (1980) model (red curve) for comparison. b: dry shear and c: bulk moduli for some chalk data. solid lines are upper (red) and lower (green) modified hashin-shtrikman boundaries. following compaction changes new moduli are estimated by shifting points parallel to these boundaries. slightly modified from vejbæk et al. (in press b). 19 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.91800 2000 2200 2400 2600 0.05 0.1 0.15 0.2 0.25 0.3 0.35 0.4 0.45 0.5 1.7 1.75 1.8 1.85 1.9 1.95 2 6 7 8 9 10 11 12 131800 2000 2200 2400 2600 1800 2000 2200 2400 2600 1800 2000 2200 2400 2600 1800 2000 2200 2400 2600 1800 2000 2200 2400 2600 1800 2000 2200 2400 2600 2000 6000 10000 14000 2000 nana-1xp m-10x m-9x 6000 top chalk base chalk 10000 14000 2000 6000 10000 14000 2000 6000 10000 14000 2000 6000 10000 14000 2000 6000 10000 14000 2000 6000 10000 14000 2000 6000 10000 14000 2600 2400 2200 2000 1800 c: porosity realisation e: vp/vs ratio realisation f: impedance realisation g: seismic realisation with oil h: seismic realisation without oil d: oil saturation b: interpreted horizonsa: field seismic fig. 3. a: original seismic section; see fig. 1 for location. b: depth profile of the interpreted horizons. the lower two thirds of the chalk is based only on the nana-1xp well. c: reconstructed porosity section. note lateral porosity variation outside well control. d: calculated water saturation profile. the free water level (green line) controls the calculated oil columns (cf. fig. 1b). e: calculated vp/vs ratio. oil causes a reduced vp/vs ratio. f: calculated acoustic impedance profile. the porosity dependency is clear, but only a weak hydrocarbon effect is present. g: forward modelled seismic response with the oil content as in d. h: forward modelled seismic response with only water. a subtle amplitude increase may be noted at top chalk level. vertical and horizontal scales are in metres. example from the dan and halfdan fields seven horizons were interpreted in the chalk in the studied case. interpolation of rock properties occurs along these horizons. the deeper two thirds of the chalk succession has only been penetrated by the nana-1xp well, whereas the upper third (and shallower horizons) have been drilled by all wells. properties are therefore only interpolated in the shallower parts, whereas deeper parts are extrapolated from the nana1xp well (fig. 3b). in addition an fwl (based on albrechtsen et al. 2001) is included to show fluid effects (fig. 1). the resulting peculiar oil saturation variations (fig. 3c) are due to the zig-zag course of the profile. depth dependent changes in porosity beyond well control are obvious and are reflected in the acoustic impedance profile (fig. 3d, e). oil effects on impedance are insignificant (fig. 3c, e), whereas a clear fluid response is seen in the vp/vs ratio profile (fig. 3f). fluid effects are thus predicted to be detectable with avo processing from which the vp/vs ratio is obtained. however, due to the non-linear fluid effect on the vp/vs ratio, it is not well suited for monitoring reservoir depletion in real data (timelapse seismic) which is affected by noise (e.g. japsen et al. 2004, 2005; vejbæk et al. in press b). the most significant vp/vs response occurs in the oil saturation interval of 0% to 30%, whereas the response change from typical residual oil saturation of around 30% to full saturation is negligible. the selective fluid predicting capability of vp/vs (or poisson ratio) is still of considerable value for reservoir mapping because chalk reservoirs generally are affected by tilted fluid contacts that may be difficult to predict (e.g. vejbæk et al. in press a). the two forward models of the reflection seismic response (fig. 3g, h) demonstrate the subtle fluid effect on seismic response. in fig. 3h all oil has been replaced by water. this causes a slight increase in top chalk reflectivity. conclusions the applied compaction model reproduces changes in seismic reflectivity similar to those observed. hydrocarbons are shown to have a subtle impact on standard seismic response in chalk. direct oil identification from reflection data is unlikely in the absence of abnormally preserved porosity by early hydrocarbon invasion. however, direct identification of hydrocarbons based on the vp/vs ratio (or poisson ratio) derived from amplitude versus offset analysis may be possible. in some cases, impedance data may show fluid effects. in general the fluid effects are subtle and confirmation by forward modelling requires models to be as realistic as possible. references albrechtsen, t., andersen, s.j., dons, t., engstrøm, f., jørgensen, o. & sørensen, f.w. 2001: halfdan: developing non-structurally trapped oil in north sea chalk. society of petroleum engineers paper 71322, 14 pp. britze, p., nielsen, e.b., dahl, n. & haug, s. 2000: north sea chalk porosity resolved by integration of seismic reflectivity and well log data. eage/said conference, paris, france, 6–8 november, 2000. abstracts. paper b30, 7 pp. engstrøm, f. 1995: a new method to normalize capillary pressure curves. international symposium of the society of core analysts, san francisco, usa, 1995. proceedings. paper 9535, 12 pp. fabricius, i.l., mavko, g., mogensen, c. & japsen, p. 2002: elastic moduli of chalk as a reflection of porosity, sorting and irreducible water saturation. society of exploration geophysicists technical programme and expanded abstracts 21, 1903–1906. gassmann, f. 1951: elastic waves through a packing of spheres. geophysics 16, 673–685. jacobsen, n.l., engstrøm, f., uldall, a. & petersen, n.w. 1999: delineation of hydrodynamic/geodynamic trapped oil in low permeability chalk. society of petroleum engineers paper 56514, 10 pp. japsen, p. 1998: regional velocity–depth anomalies, north sea chalk: a record of overpressure and neogene uplift and erosion. american association of petroleum geologists bulletin 82, 2031–2074. japsen, p., høier, c., rasmussen, k.b., fabricius, i.l., mavko, g. & pedersen, j.m. 2002: effects of fluid substitution on ultrasonic velocities in chalk plugs, south arne field, north sea. society of exploration geophysicists technical programme and expanded abstracts 21, 1881–1884. japsen, p., bruun, a., fabricius, i.l., rasmussen, r., vejbæk, o.v., pedersen, j.m., mavko, g., mogensen, c. & høier, c. 2004: influence of porosity and pore fluid on acoustic properties of chalk: avo-response from oil, south arne field, north sea. petroleum geoscience 10, 319–330. japsen, p., bruun, a., fabricius, i.l. & mavko, g. 2005: identification of hydrocarbons in chalk reservoirs from surface seismic data; south arne field, north sea. geological survey of denmark and greenland bulletin 7, xx–yy (this volume). jensen, p.k., holm, l. & thomsen, e. 1985: modelling burial history, temperature and maturation. in: thomas, b.m. et al. (eds): petroleum geochemistry in exploration of the norwegian shelf, 145–152. london: graham & trotman. sclater, j.g. & christie, p.a.f. 1980: continental stretching: an explanation of the post-mid-cretaceous subsidence of the central north sea basin. journal of geophysical research 85, 3711–3739. vejbæk, o.v., frykman, p., bech, n. & nielsen, c.m. in press a: the history of hydrocarbon filling of danish chalk fields. in: doré, a.g. & vinning, b. (eds): petroleum geology: north-west europe and global perspectives: proceedings of the 6th petroleum geology conference. london: geological society. vejbæk, o.v., rasmussen, r., japsen, p., bruun, a., pedersen, j.m., marsden, g. & fabricius, i.l. in press b: modelling seismic response from north sea chalk reservoirs resulting from changes in burial depth and fluid saturation. in: doré, a.g. & vinning, b. (eds): petroleum geology: north-west europe and global perspectives: proceedings of the 6th petroleum geology conference. london: geological society. walls, j.d., dvorkin, j. & smith, b.a. 1998: modeling seismic velocity in ekofisk chalk. society of exploration geophysicists technical programme and expanded abstracts 17, 1016–1019. 20 authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ov@geus.dk geological survey of denmark and greenland bulletin 26, 2012, 25-28 25 natura 2000 habitat mapping in kattegat, denmark: an example from læsø trindel zyad k. al-hamdani and laura g. addington natura 2000 is a network of nature protection areas established by the european union under the habitats directive (european union 1992). the aim is to assure long-term survival of the most valuable and endangered species and habitats in europe. the network comprises special areas of conservation and protection designated by the member states under, respectively, the habitats directive and the birds directive. the establishment of the network of protected areas also fulfils a community obligation under the convention of biological diversity of the united nations. the member states of the european union are obliged to ensure full compliance with natura 2000 legislation and must work out a management plan for each natura 2000 area. in the marine environment, detailed information on the natura 2000 areas is necessary to make informed management plans in fulfilment of national and international obligations and regulate human activities. the data may come from geophysical mapping or from biological sampling or from both. this paper presents some of the results of a habitat mapping project performed in 2011 covering almost all natura 2000 areas and pre-designated, aggregate extraction areas in the inner danish waters. due to time and resource limitations, the natura 2000 areas in danish waters were mapped by geophysical methods to produce seabed sediment maps, followed by biological sampling at selected stations. this is a fast and reliable method for producing a broad-scale habitat map of large areas in a limited time. the areas were mapped and classified according to the habitats directive (european union 1992; european commission 2007). the mapping was also conducted to conform to the marine strategy framework directive, which requires the member states to provide an initial assessment of their marine waters by july 2012 (european union 2008). the geological survey of denmark and greenland (geus) carried out the geological and geophysical work, and orbicon a/s undertook the biological investigations in 18 natura 2000 areas in kattegat and the southern baltic sea (fig. 1). læsø trindel is a natura 2000 area with a significant biodiversity and a wide range of habitat types of which three are of special interest: (1) boulder reefs, (2) sandbanks and (3) ‘bubbling reefs’ which are structures formed by leaking gas. geological setting læsø trindel is located north-east of læsø, the largest island in kattegat. the læsø trindel area is located in the nw– se-trending fennoscandian border zone and has been affected by neotectonic deformation (hansen 1995). in this area there is a succession of weichselian clay and silt with scattered stones and boulders, locally overlain by till. after isostatic rebound and erosion of the sediments in the holocene, a widespread residual layer of boulders or patches of boulders has been left. hermansen & jensen (2000) produced a regional (1:500 000 scale) seabed sediment map of the danish waters, which shows the presence of till deposits in the area on and around læsø trindel. however, at that time there were insufficient data for a detailed sediment map of the area. 10°e 12°e 14°e 2 km trindel kattegat baltic sea læsø sailing lines: 2005 project sailing lines: current project aggregate extraction areas natura 2000 areas 56°n denmark læsø 50 km 11.2° 57.44° fig. 1. map of inner waters in denmark showing the locations of the natura 2000 and the aggregate extraction areas mapped in 2011. inset: 2005 and 2011 sailing lines in the læsø trindel area. © 2012 geus. geological survey of denmark and greenland bulletin 26, 25–28. open access: www.geus.dk/publications/bull 2626 a more recent study undertaken by geus in 2005 yielded echo sounder, side-scan sonar and ground truth data from the central part of læsø trindel (fig. 1; leth & al-hamdani 2008) showing that it consists of a 2 × 2 km shallow platform with water depths of 4 to 13 m. in the eastern and south-eastern parts of the survey area, the water depth increases to 45 m (fig. 2a). the platform consists mainly of till with boulders and cobbles that vary in size and distribution throughout the area. its central part shows the highest density of cobbles and boulders, and this is the main boulder reef of the region. the surrounding part is characterised by gravelly sand and sand patches with scattered boulders. methods field work in 2011 was carried out in two phases: in phase one, bathymetry data were collected with a single-beam echo sounder, and overlapping seabed backscatter images were obtained using side-scan sonar with 100 and 400 khz frequencies. information on the sub-bottom layers was obtained with a chirp profiler (1–10 khz; providing high resolution but low penetration) and a sparker (1 khz; providing deeper penetration but lower resolution). the area was surveyed with 100 m line spacing to ensure full seabed coverage. cross lines with 4 km spacing were surveyed for cross-verification (fig. 1). in phase two, which succeeded preliminary interpretation of the acoustic data, areas of special interest were video filmed with a camera mounted on a remotely operated vehicle, and surface sediment samples were collected with a van veen grab sampler. finally, sediment cores were collected with a 6 m long vibrocorer. the acoustic data, the video films and the grab and core samples were used together with older data for the final interpretation and habitat mapping . results and discussion the data from 2005 and 2011 cover almost the entire læsø trindel natura 2000 area (fig. 1) and were used to map th e bathymetry along the sailing lines, to compile a side-scan mosaic and to map the distribution of various substrates and nature types. the depth data from 2011 compare well with the depth data from 2005 which were acquired with a highly accurate multibeam echo sounder (fig. 2a). the side-scan backscatter mosaic image shows variation in substrate reflectivity due to the different signatures of the different sediments (fig. 2b). the interpretation of the image was carried out by dividing the image into 50 × 50 m blocks and manually classifying the sediments in each block. fig. 2. the læsø trindel natura 2000 area. a: bathymetry. b: side-scan sonar mosaic. the grey scale reflects hard (dark) to soft (light) sediments. c: interpreted distribution of substrate types and natura 2000 habitats. ’bubbling reef’: current project ’bubbling reef’: 2005 project substrate type 4 substrate type 3 substrate type 2 substrate type 1b substrate type 1a natura 2000 area water depth (m) 0 6 8 10 13 16 19 22 25 28 31 40 50 2 km a b c 11.2°e 57.44°n 27 interpretation of the side-scan images together with ground truth samples allowed us to produce a map that shows the distribution of substrate types and the occurrence of natura 2000 habitat types, using the classification categories suggested by the danish nature agency (table 1; fig. 2c). the chirp profiles were also used for the interpretation. the different types of sediments give different seismic signals that enhance the final seabed sediment interpretation significantly. boulder reefs (substrate type 4) cover c. 38% of the na tura 2000 area. between these, substrate type 2 with sand and scattered boulders dominates. the deeper south-eastern parts of the surveyed area are dominated by soft sediment. sandbanks are not well defined in the læsø trindel area, but substrate description type 1a soft bottom sediment with silt and mud 1b sand, including sandbanks 2 sand, gravel and small stones <10 cm. may also comprise scattered stones >10 cm that cover <10% of the area 3 sand, gravel and small stones <10 cm. may also comprise scattered stones >10 cm that cover 10–25% of the area 4 stones >10 cm with >25% coverage; also with sand, gravel and scattered, small stones table 1. substrate classification fig. 3. examples of ‘bubbling reefs’. a: on side-scan images ‘bubbling reefs’ appear as small areas with lowered reflectivity (arrows). b: shallow seismic sparker profiles with sites where the water column is disturbed by gas seeping out from the sediments (arrows). c: snapshots from a video film taken with a camera mounted on a remotely operated vehicle. gas bubbles are seen in the water column above the ‘bubbling reefs’ (arrows). the reefs have a rich epiflora and epifauna. a b c 25 m 25 m 7.5 m100 m 7.5 m100 m 2828 some sand accumulations (substrate type 1b) are found in the eastern and southern parts. ‘bubbling reefs’ form at sites where methane gas seeping up from eemian and early weichselian marine deposits has led to carbonate cementation of sand in the methane oxidation zone (jensen et al. 1992; laier et al. 1992). in some areas, cemented sand slabs have been exposed by subsequent erosion producing a hard substrate, which can be colonised by plants and animals that form a rich epiflora and epifauna. the geology, geochemistry and biology of the ‘bubbling reefs’ have been intensively studied, and their structure, origin and formation have been described by jørgensen (1989), jensen et al. (1992) and laier & jensen (2007). ‘bubbling reefs’ are fairly common in the læsø trindel region where they were discovered by chance by fishermen and sport divers. some ‘bubbling reefs’ were recorded on sidescan images during previous surveys (leth & al-hamdani 2008), and additional occurrences of these structures were recorded in the side-scan data from 2011 (fig. 2c). they are distinctive on the side-scan images (fig. 3a), but can also be detected on shallow seismic profiles (fig. 3b) where gasprone structures can be mapped, and sometimes gas bubbles can be seen in the water column. the ‘bubbling reefs’ occur mainly in areas with sandy seabed (substrate type 1b). the reefs support benthic communities rich in species (fig. 3c) and have many cavities where animals can find shelter. concluding remarks 1. habitat mapping using geophysical methods in combination with ground truth observations and sampling is a cost-eff ective means of mapping large seabed areas in a relatively short time. 2. some natura 2000 habitats such as the ‘bubbling reefs’ were originally discovered by chance by fi shermen and sport divers. side-scan surveying is an eff ective way to map this habitat type over larger areas. 3. th e results obtained during this project provide information on the physical features and the habitat types of læsø trindel. th ese two elements are important to assess the current environmental status of marine waters as required by the marine strategy framework directive (table 1 in annex iii). 4. management of human activities in the marine environment requires a management plan that ensures habitat protection and at the same time enhances sustainable economic growth of the region. th is project provides information on natura 2000 areas for decision makers in denmark. acknowledgements the danish nature agency is thanked for financing the project and for permission to publish the results. we also thank orbicon a/s for ground truth samples and films, and the captain and the crew of r/v laura for help during the cruises. references european commission 2007: interpretation manual of european union habitat, 142 pp. brussels: european commission, http://ec.europa.eu/ environment/nature/legislation/habitatsdirective/docs/2007_07_im.pdf european union 1992: council directive 92/43/eec of 21 may 1992 on the conservation of natural habitats and of wild fauna and flora (habitat directive). official journal of the european union l 206, 7–92, http://eur-lex.europa.eu/lexuriserv/lexuriserv.do?uri=oj:l:1992: 206:0007:0050:en:pdf european union 2008: directive 2008/56/ec of the european parliament and of the council establishing a framework of community action in the field of marine environment policy (marine strategy framework directive). official journal of the european union l 164, 19–40, http://eur-lex.europa.eu/lexuriserv/lexuriserv.do?uri=oj:l:2008: 164:0019:0040:en:pdf hansen, j.m. 1995: en ø’s opståen, kystdannelse og vegetationsudvikling: naturlige og menneskeskabte landskaber på læsø. geologisk tidsskrift 2, 1–74. hermansen, b. & jensen, j.b. 2000: digitalt kort over havbundssedimenter omkring danmark 1:500.000. danmarks og grønlands geologiske undersøgelse rapport 2000/68 (cd-rom). jensen, p., aagaard, i., burke jr., r.a., dando, p.r., jørgensen, n.o., kuijpers, a., laier, t., o´hara, s.c.m. & schmaljohann, r. 1992: ‘bubbling reefs’ in kattegat: submarine landscapes of carbonate-cemented rocks support a diverse ecosystem at methane seeps. marine ecology progress series 83, 103–112. jørgensen, n.o. 1989: holocene methane-derived, dolomite cemented sandstone pillars from the kattegat, denmark. marine geology 88, 71–81. laier, t. & jensen, j.b. 2007: shallow gas depth-contour of skagerrak – western baltic sea region. geo-marine letters 27, 127–141. laier, t., jørgensen, n.o., buchardt, b., cederberg, t. & kuijpers, a. 1992: accumulation and seepages biogenic gas in northern denmark. continental shelf research 12, 1173–1186. leth, j.o. & al-hamdani, z. 2008: mapping of the natura 2000 annex 1 habitats 1170 and 1180 in n kattegat, combining acoustic and ground truth methods. in: dinesen, g.e. (ed.): mapping and modelling of marine habitats in the baltic sea region. balance interim report 27, 75–89. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: azk@geus.dk geological survey of denmark and greenland bulletin 26, 2012, 45-48 45 results of monitoring groundwater above the natural gas underground storage at stenlille, denmark troels laier groundwater in the stenlille area is regularly analysed for light hydrocarbons after a natural gas underground storage facility was established there in 1989. the monitoring is carried out by the geological survey of denmark and greenland and is part of the authorities’ requirements for the environmental approval of the natural gas storage run by the state-owned danish oil and gas company dong a/s. groundwater from observation wells and water wells in the area was analysed every month during the first year of operation and four times a year in the following years. more frequent analyses are undertaken on special occasions. underground gas storage two underground storage facilities were established in denmark in order to buffer the supply of gas from the north sea, one in salt caverns in jylland, the other in a deep aquifer at a depth of 1500 m near stenlille (fig. 1). natural gas underground storage is possible in deep sealed geological formations with good reservoir quality and covered by tight caprock. these conditions are met by an anticlinal structure near stenlille with a vertical closure of c. 35 m covering an area of 14 km2. the gassum formation forms the reservoir where gas is stored by displacing formation water. the formation is of late triassic age and consits of cyclically interbedded sandstone and marine mudstone that were deposited as a result of changes in the depositional environment (hamberg & nielsen 2000). the overlying 300 m thick lower jurassic fjerritslev formation, which consists of claystone, serves as a caprock for the sandstone reservoir (fig. 2). the total estimated storage capacity of the stenlille structure equals three billion normal cubic metres, and due to reservoir heterogeneities, gas is stored in several separate zones (fig. 1). the gas storage is operated by 14 wells for injection and withdrawal of gas and six wells used for observational purposes, most of them in the periphery of the structure. fig. 1. map of the stenlille area showing the location of water wells (blue), observation wells for shallow aquifers (green, inset map) and deep wells (red). the extent of the underground gas storage is indicated by coloured areas, where yellow and red show different gas zones. stenlille 11°35´e 55°33´n 55°33´n 100 m k1 k2 1 km st14 st2 529 400 558558 407 257 502 520 nyrup 518 stenlille © 2012 geus. geological survey of denmark and greenland bulletin 26, 45–48. open access: www.geus.dk/publications/bull 4646 baseline study as part of the pre-investigation of the stenlille stucture, a baseline study was carried out in order to characterise the type and concentration of natural hydrocarbons present in the rocks within and above the future gas storage reservoir. analyses of hundreds of drill cuttings from deep wells indicated no significant occurrence of hydrocarbons, although low concentrations of in situ generated hydrocarbons appeared to be present in the organic rich fjerritslev formation (laier & øbro 2009). a very low concentration of methane (30 mg/l), possibly of microbial origin, was observed in the saline formation water of the gassum formation (laier & øbro 2009). bacterial methane in groundwater methane is generally not analysed in groundwater unless its presence is suspected; therefore very few analyses existed prior to the baseline study. groundwater from 21 water wells in the stenlille area was sampled, and methane was found to occur in low concentrations (0.01–0.49 mg/l) in all of them except two (laier 1989). no higher hydrocarbons were detected in any of the samples. stable isotope analyses of nine of the samples showed that the methane was of bacterial origin (δ13cch4: –90 to –62‰; laier 1989). the bacterial methane most likely comes from peat layers in the area. most water wells in the area draw water from glacial meltwater sand found at depths of 20–30 m and overlain and underlain by glacial till (fig. 3). bacterial methane (0.40 mg/l; δ13c1 = –64.5‰) was also found in groundwater from the shallow filter of the k1 observation well located adjacent to the first injection well (st2; fig. 1). the methane concentration in the brackish water from the deep filter (fig. 3) was too low (0.02 mg/l) to permit analysis of its carbon isotope ratio. natural gas storage storage of natural gas from the danish part of the north sea began in july 1989. the gas is dominated by methane (91%) with a little ethane (5.5%) and propane (2.0%). the average δ13cch4 value of methane is –46.6‰. the composition of the natural gas stored in the stenlille structure is markedly different from that of the gas dissolved in shallow groundwater (fig. 4). this makes it fairly easy to distinguish between the two types of gas and identify even minor leaks from the deep gas storage reservoir at the near surface. groundwater methane a fairly large variation in the concentration of methane from 0.4 to 3.7 mg/l was observed in water samples from a shallow filter of the k1 observation well in the early 1990s, but not in samples from a deeper filter, which had very low values (0.02 mg/l; fig. 5). if the variation in methane concentraclaystone limestone clay, sand, till zechstein salt sandstone 0 500 1000 1500 well for injection and production observation well dep th (m ) gas fig. 2. sketch of the natural gas underground storage facility at stenlille. a zechstein salt pillow is located below the surface. the development of the salt pillow has deformed the overlying beds into a dome. sandstone: gassum formation sandstone (late triassic), claystone: fjerritslev formation (early jurassic). fig. 3. geological cross section of the upper layers of the stenlille natural gas underground storage. the locations of wells k1, k2 and 558 are shown on fig. 1. the red colour indicates the distribution of gas after the st14 leakage in august 1995 (for details see laier 2010). ? ?? ? elevation (m) –80 –60 –40 –20 0 20 40 enewsw 558 paleocene calcareous sand paleocene marine clay glacial till meltwater sand 500 m gas after leakage in august 1995 k2 k1 ??? 47 tion was due to a leakage around the nearby st2 well (fig. 1), the opposite might be expected with an increase in methane at the deeper level first. in order to deduce the origin of the methane, isotope analyses were regularly performed during the first year of sampling. the isotope ratio of the methane (δ13cch4: –62 to –52‰) together with a lack of higher hydrocarbons indicate that the methane was bacterial in origin (fig. 4), though its source remains unknown. since the methane concentration was highest during the first years after drilling the k1 well, one may speculate whether hydrogen generated by anaerobic corrosion of iron filings from drilling may have led to methane formation as suggested by daniels et al. (1987). groundwater methane concentration in water from the stenlille waterworks and from private water wells remained at the low levels measured prior to the storage of natural gas (fig. 6). minor gas leakage in september 1995, gas bubbles were observed at the terrain surface near the newly drilled st14 injection well (fig. 1), and a sample of the gas was immediately collected and analysed. its chemical and isotopic composition was similar to that of the gas being stored (fig. 4), so there was little doubt that gas leaked from the new well. the leak was due to technical problems during gas injection (laier & øbro 2009) and was quickly stopped, but an estimated 5000 m3 of gas was lost to geological formations above the reservoir cap rock. a week after the leak was observed at the surface, a significant increase in the dissolved gas concentration was measured in the deep filter of the k1 well, located 250 m from the st14 well. this gas had almost the same composition as the reservoir gas (fig. 4). no free gas was observed during collection of water from this level, so it was concluded that all the gas was dissolved at this time. however, for the gas to migrate from st14 to k1 during a week, a free gas phase –100 c 1 /(c 2 +c 3 ) 105 103 102 10 0 104 –80 –60 –40 –20 δ13c ch4 (‰) genic gas thermomicrobial gas dissolved gas in the upper aquifer of the k1 observation well dissolved gas in the shallow aquifers natural gas from the danish north sea dissolved gas in the lower aquifer of k1 from the gas leakage in 1995 fig. 4. classification diagram showing the chemical and isotopic composition of natural gas from the danish north sea and dissolved gas in shallow aquifers. the c1/(c2+c3) ratio for the dissolved gas represents a minimum value as both ethane and propane were below detection limit for all dissolved gases. the white areas are typical of thermogenic and microbial gas and the grey area represents either mixtures of these gas types or gas of unknown origin. c1: methane concentration, c2: ethane concentration, c3: propane concentration. 10 year 89 91 93 95 97 99 01 03 05 07 09 11 c h 4 ( m g/ l) 0 2 4 6 8 10 start of injection 98–128 m 35–45 m 27 mg/l fig 5. methane concentration in groundwater from the k1 observation well. the sudden increase in dissolved methane was noted one week after the gas leakage at st14. c h 4 ( m g/ l) 0 0.1 0.2 0 0.1 0.2 c h 4 ( m g/ l) 0.3 0.4 year 89 91 93 95 97 99 01 03 05 07 09 11 520 257 502 529 518 400 558 407 a b fig. 6. methane concentration in groundwater from wells supplying the stenlille waterworks (a) and from private wells (b). 4848 must have existed at some point, although free gas was never observed in the samples from any of the wells. the concentration of dissolved gas has decreased since the gas leakage occurred and is presently below 1 mg/l, and its composition has hardly changed. traces of natural gas in a water well near nyrup the regular groundwater survey revealed a significant increase in methane in october 2009, from 0.1 to over 0.4 mg/l, in water well 558 south-west of nyrup (figs 1, 6b). traces of ethane and propane were also observed, indicating the presence of natural gas in low concentrations. an immediate follow-up investigation indicated a high proportion of old groundwater in the well due to a low pumping rate. when the normal pumping rate for the well was reestablished, groundwater methane dropped below 0.1 mg/l and the higher hydrocarbons ethane and propane could no longer be detected. it was concluded that minor amounts of natural gas from the leak in 1995 had migrated into the paleocene calcareous sand below paleocene marine clay of low permeability and reached the water well near nyrup (fig. 3; laier 2010). a buried valley at nyrup may have allowed gas to migrate to shallower depths where it gradually dissolved in groundwater (jørgensen & sandersen 2009). the very low pumping rate in well 558 resulted in less infiltration of younger groundwater, which enabled the detection of natural gas in older groundwater (laier 2010). conclusions no leakage from the natural gas underground storage has occurred through the natural (geological) barriers. regular analyses of dissolved hydrocarbons in shallow groundwater only showed the presence of bacterial methane, which was also present prior to the storage of natural gas. the traces of natural gas observed for a short period in one of the water wells in 2009 could be related to a minor leak from a newly drilled injection well in august 1995. references daniels, l., belay, n., rajagopal, b.s. & weimer, p.j. 1987: bacterial methanogenesis and growth from co2 with elemental iron as the sole source of electrons. science 237, 509–511. hamberg, l. & nielsen, l.h. 2000: shingled, sharp-based shoreface sandstones: depositional response to stepwise forced regression in a shallow basin, upper triassic gassum formation, denmark. in: hunt, d. & gawthorpe, r.l. (eds): sedimentary reponses to forced regressions. geological society special publications (london) 172, 69–89. jørgensen, f. & sandersen, p. 2009: kortlægning af begravede dale i danmark. opdatering 2007–2009, 374 pp. copenhagen: geological survey of denmark and greenland. http://www.begravededale.dk/ pdf_2009/rapport_begravede_dale_2009_137mb.pdf laier, t. 1989: stenlille gas storage – study of naturally occurring hydrocarbon gases before injection. dgu service report 15, 53 pp. copenhagen: geological survey of denmark. laier, t. 2010: spor af naturgas i nyrup vandværksboring oktober 2009. danmarks og grønlands geologiske undersøgelse rapport 2010/1, 30 pp. laier, t. & øbro, h. 2009: environmental and safety monitoring of the underground gas storage facility at stenlille, denmark. in: evans, d.j. & chadwick, r.a. (eds): underground gas storage: worldwide experiences and future development in the uk and europe. geological society special publications (london) 313, 81–92. author’s address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: tl@geus.dk geological survey of denmark and greenland bulletin 23, 2011, 25–28 25 a major hydrogeological programme has been carried out to map the miocene succession in central and southern jylland (fig. 1). the miocene deposits comprise several aquifers with potential drinking water resources and have been investigated by drilling and acquisition of seismic data integrated with sedimentology and biostratigraphy. scharling et al. (2009) described a 3d hydrogeological model that covers part of the onshore danish miocene deposits. the model was based on a sequence-stratigraphic approach and led to a better understanding of the geological architecture of the aquifers than traditional lithofacies models. hence it was decided to establish a digital, spatial, geological model covering the entire onshore miocene succession (kristensen et al. 2010). geology the onset of the miocene is characterised by inversion tectonics causing a change in the depositional regime from full marine, clayey sediments to shallow-water, sand-rich, delta deposits (rasmussen et al. 2010). during the early – early middle miocene, regressions and transgressions were strongly controlled by eustatic sea-level changes, resulting in three phases of shoreline progradation into the basin that covers present-day denmark. the three phases are represented by the sand-rich billund, bastrup and odderup formations, intercalated with the clayey and silty marine vejle fjord, klintinghoved and arnum formations (fig. 2). during the late middle miocene to the late miocene, the marine clay of the hodde, ørnhøj and gram formations were deposited and towards the end of the miocene a new progradation resulted in deposition of the sandy marbæk formation. sequence stratigraphic framework the miocene digital, spatial, geological model is based on the sequence-stratigraphic framework of rasmussen (2004) and rasmussen & dybkjær (2005) and the lithostratigraphy of rasmussen et al. (2010). these studies are based on new borehole data, high-resolution seismic profiles (vangkildepedersen et al. 2006) and high-resolution biostratigraphy (dybkjær & piasecki 2010). in sequence stratigraphy a geological succession is divided into a succession of different lithofacies (a sequence) bounded by key-surfaces and commonly stacked in a cyclic manner. each sequence represents a digital, spatial, geological model of the miocene in jylland, denmark margrethe kristensen, thomas vangkilde-pedersen and erik skovbjerg rasmussen 10°e 14°e 55°n 57°n sweden denmark jylland fig. 5 100 km fig. 1. map of denmark showing the study area in central and southern jylland (rectangle). m io ce n e 10 15 20 u p p er m id d le l o w er swepochma lithostratigraphy ne marine silt and clay marine sand fluvial sand and gravel hiatus brackish-water silt and clay coal marbæk fm gram fm ørnhøj fm hodde fm arnum fm odderup fm bastrup fm resen mb klintinghoved fm vejle fjord fm brejning fm brejning fm billund fm addit mb fig. 2. lithostratigraphic scheme of the miocene of onshore denmark showing the distribution of the formations from south-west to north-east jylland. marine deposits dominate in the south-western part of jylland. delta or fluvial sand deposits are mainly found in the central parts of jylland, where they form large potential groundwater reservoirs (rasmussen et al. 2010). © geus, 2011. geological survey of denmark and greenland bulletin 23, 25–28. open access: www.geus.dk/publications/bull 2626 a cycle in relative sea level and can be subdivided into four systems tracts that link contemporaneous deposits together. (1) during the early stage of sea-level rise the lowstand systems tract is formed when the sediment supply from the hinterland is greater than the sea-level rise. the shoreline progrades into the basin and sands and clays are deposited in association with delta progradation. incised valleys are filled up by fluvial deposits, which are usually dominated by coarse-grained sediments. (2) the transgressive systems tract is formed when the base-level rise outpaces the sediment supply from the hinterland, causing the shoreline to move landwards. this landward movement of the shoreline results in predominant clay and silt deposition on the former delta platform and in the basin. however, sand is still deposited along the shoreline, commonly on the shoreface and in inlets of barrier complexes, as bars in tidally influenced estuaries and in incised valleys associated with fluvial systems. (3) during the late stage of sea-level rise the highstand systems tract is formed. the sediment supply from the hinterland outpaces the sea-level rise and the shoreline again progrades into the basin. the highstand systems tract commonly shows fineto coarser-grained deposits laid down on the slope of delta complexes or as shoreface sands alternating with lagoonal clays. (4) the forced regressive wedge systems tract is formed in the marine part during falling sea level and is hence characterised by progradation of the shoreline. during falling sea level, incision commences on the highstand delta complex. the deposits are typically dominated by well-sorted, relatively coarse-grained sediments. common deposits include different types of off-lapping shallow marine shoreface and delta deposits and deep-sea submarine fans. in the miocene succession of denmark some of the best aquifers are associated with lowstand systems tracts (fig. 3). the sand that constitutes the aquifers was partly deposited in incised valleys and partly as prograding deltas in the basinal area. the incised valleys are dominated by fluvial deposits, namely the addit member of the billund formation and the resen member of the bastrup formation (fig. 2). both generations show a two-fold subdivision of the valley fill that coincides with delta progradation into the basin. this pattern is, however, interrupted by minor flooding where a thin sequence of marine sand, clay and coal was deposited. an example of two such successive delta deposits from the lowstand systems tract of the bastrup formation is shown in fig. 4. the glacio-eustatic sea-level changes in the miocene resulted in an asymmetric pattern of slow regressions and rapid transgressions, which explains why transgressive sand or clay was rarely deposited. modelling of the miocene succession the backbone of the miocene digital, spatial, geological model is nine correlation panels (five w–eand four s–noriented), which constitute a conceptual geological model of the danish miocene (rasmussen et al. 2010). the conceptual model is based on sedimentological investigations of samples from 150 boreholes (c. 100–400 m deep), detailed biostratigraphical studies of samples from 50 boreholes and studies of 25 outcrops tied together with a dataset of c. 1200 km high-resolution seismic profiles. with the conceptual model as the starting point, a 3d geological model has been established using the software package geoscene 3d (www.i-gis.dk). the geoscene 3d software gives the user direct access to carry out interpretation moving through the subsurface and better understand 3d structures. borehole data, geophysical logs and seismic coastal plain coastal marine lagoon prodelta/offshore incised valley fill sequence boundary subaerial unconformity and correlative surface lst2 lst1 frwst hst lst2 lst1 fig. 3. a conceptual model for the development of the miocene deposits in jylland. hst: highstand systems tract, frwst: forced regressive wedge systems tract. lst2: lowstand systems tract unit 2. lst1: lowstand systems tract unit 1. top miocene lst1 lst2 base miocene top bastrup 125 m 50 m fig. 4. seismic profile showing two successive delta complex deposits of the lowstand systems tract of the bastrup formation. lst1: lowstand systems tract unit 1. lst2: lowstand systems tract unit 2. for location see fig. 5b. 27 data have been imported to the modelling software and interpretation performed both on 2d profiles and in the 3d environment. the set-up is constructed as a layer model, but emphasis has been on distinguishing between different generations of delta lobes as the shoreline prograded into the basin during deposition of the billund, bastrup and odderup formations. thus the model comprises 75 layers and lithological units, which have been named according to the formation and depositional environment. the top of each lithological unit is interpreted using interpretation points. the interpretation is based on data from: stratigraphically described boreholes, high-resolution seismic profiles or boreholes from the national borehole database. free digital points have also been added to indicate the outline of deltalobes. the interpretation includes an evaluation of the quality of the points. the top miocene, top bastrup, top billund and base miocene surfaces are interpreted in almost the entire model area. unlike these surfaces, the propagation of each generation of delta lobes, is limited and follows the position of the coastline, at the time of deposition. on the basis of high-resolution seismic profiles and detailed lithological descriptions of borehole samples, the extension of each delta lobe has been interpreted. in fig. 5 the maximum extension of 10 generations of delta lobes of the bastrup formation and 11 generations of delta lobes of the billund formation is shown together with the coverage of high-resolution seismic profiles and boreholes used in the sequence-stratigraphic interpretation. to support the model, detailed interpretation of the seismic profiles shown in fig. 5 has been conducted with focus on mapping the extent of sand-rich bodies. all previous interpretations of the top miocene, top billund, top ba-strup and base miocene have been checked, and if necessary, revised, according to the present level of knowledge. top and bottom of sand-rich bodies in the form of delta or fluvial deposits have been interpreted on all seismic lines in order to assist the modelling work in geoscene 3d. internal, parallel, clinoform, reflection patterns with dips of 5–10° have been interpreted as a direct indicator of fineto coarse-grained sand, whereas sigmoidal clinoform internal reflection patterns typically indicate alternating layers of clay and sand (rasmussen et al. 2007: bassetti et al. 2008; hansen & rasmussen 2008). incised valleys and fluvial channels expressed by concave-up erosion surfaces are typically filled with coarseto fine-grained sand and coarse-grained sand and gravel, respectively (rasmussen et al. 2007). in places where the seismic data do not directly indicate sandy deposits, the profiles have nevertheless been used to extrapolate available borehole information in the best possible way. gamma-ray logs have been a valuable supplement to the geological descriptions of borehole samples. most of the boreholes were drilled using the airlift drilling technique which may result in poor recovery of coarse silt and fine sand (ditlefsen et al. 2008). therefore, gamma-ray logs have been useful both for checking and correcting the lithological logs and as an indicator of depositional environment. sand-rich delta units are generally coarsening upwards and are seen on the logs as upward-decreasing gamma-ray values. fluvial channel deposits are characterised by fining-upward trends and are seen as upward-increasing gamma-ray values (fig. 6). fyn 55°n 56°n 9°e 9°e jylland fyn 55°n 56°n 9°e 9°e jylland a b fig. 6 fig. 4 seismic profile borehole delta lobe 20 km fig. 5. map of southern and central jylland showing the maximum extent of individual generations of delta lobes of a: billund formation and b: bastrup formation. the interpretation is based on high-resolution seismic profiles and data from boreholes. 2828 future work and perspectives the miocene 3d model reflects the basin development and the depositional processes as well as the palaeogeographical development during the miocene in denmark. the spatial, geological model is intended to serve as a geological database of lithological and stratigrafical information and can be seen as a visual archive of the geological knowledge of the miocene in the model area. as such, it will serve as the foundation for different types of application-oriented models. the miocene 3d model is already used in the danish nature agency as a framework for at least 11 modelling projects. the model will be updated on an annual basis in the coming years as new boreholes, seismic profiles and other data become available, or following new interpretations. acknowledgements the nature agency centres in ribe, ringkøbing and aarhus are thanked for financial support. references bassetti, m.a., berne, s., jouet, g., taviani, m., dennielou, b., flores, j.a., gaillot, a., gelfort, r., lafuerza, s. & sultan, n. 2008: the 100-ka and rapid sea level changes recorded by prograding shelf sand bodies in the gulf of lions (western mediterranean sea). geochemistry geophysics geosystems 9, q11r05, 27 pp. ditlefsen, c., sørensen, j., pallesen, t.m., pedersen, d., nielsen, o.b., christiansen, c., hansen, b. & gravesen, p. 2008: jordprøver fra grundvandsboringer, vejledning i udtagning, beskrivelse og geologisk tolkning i felten, 108 pp. geo-vejledning 1. københavn: de nationale geologiske undersøgelser for danmark og grønland. dybkjær, k. & piasecki, s. 2010: neogene dinocyst zonation for the eastern north sea basin, denmark. review of palaeobotany and palynology 161, 1–29. hansen, j.p.v. & rasmussen, e.s. 2008: structural, sedimentologic, and sea-level controls on sand distribution in a steep-clinoform asymmetric wave-influenced delta: miocene billund sand, eastern danish north sea and jylland. journal of sedimentary research 78, 130–146. kristensen, m., vangkilde-pedersen, t. & rasmussen, e.s. 2010: miocæn 3d. den rumlige geologiske model. danmarks og grønlands geologiske undersøgelse rapport 2010/91, 46 pp. rasmussen, e.s. 2004: stratigraphy and depositional evolution of the uppermost oligocene – miocene succession in western denmark. bulletin of the geological society of denmark 51, 89–109. rasmussen, e.s. & dybkjær, k. 2005: sequence stratigraphy of the upper oligocene – lower miocene of eastern jylland, denmark: role of structural relief and variable sediment supply in controlling sequence development. sedimentology 52, 25–63. rasmussen, e.s., vangkilde-pedersen, t. & scharling, p.b. 2007: prediction of reservoir sand in miocene deltaic deposits in denmark based on high-resolution seismic data. geological survey of denmark and greenland bulletin 13, 17–20. rasmussen, e.s., dybkjær, k. & piasecki, s. 2010: lithostratigraphy of the upper oligocene – miocene succession of denmark. geological survey of denmark and greenland bulletin 22, 92 pp. scharling, p.b., rasmussen, e.s., sonnenborg, t.o., engesgaard, p. & hinsby, k. 2009: three-dimensional regional-scale hydrostratigraphic modeling based on sequence stratigraphic methods: a case study of the miocene succession in denmark. hydrogeology journal 17, 1913–1933. vangkilde-pedersen, t., dahl, j.f. & ringgaard, j. 2006: five years of experience with landstreamer vibroseis and comparison with conventional seismic data acquisition. proceedings of the 19th annual sageep symposium on the application of geophysics to engineering and environmental problems, seattle, usa, 1086–1093. fluvial deposit fluvial deposit delta deposit depth (m) 70 80 90 100 110 120 top bastrup formation top bastrup fluvial sand top klintinghoved clay top bastrup delta sand micaceous clay micaceous sand quartz sand fig. 6. example of interpretation of depositional environment from gamma-ray log patterns. upward-increasing gamma-ray values are interpreted as a fining-upward f luvial deposit and upward-decreasing gamma-ray values are interpreted as an upward-coarsening delta complex. for location see fig. 5a. authors’ address geological survey of denmark and greenland, lyseng allé 1, dk-8270 højbjerg, denmark. e-mail: mkr@geus.dk geological survey of denmark and greenland bulletin 10 g e o l o g i c a l s u r v e y o f d e n m a r k a n d g r e e n l a n d b u l l e t i n 1 0 • 2 0 0 6 review of survey activities 2005 edited by martin sønderholm and a.k. higgins geological survey of denmark and greenland danish ministry of the environment geological survey of denmark and greenland bulletin 10 keywords geological survey of denmark and greenland, survey organisations, current research, denmark, greenland. cover photographs from left to right 1. the barrier reef complex fringing most of the kenya coastline is also one of the most vulnerable environments to pollution. a new oil spill sensitivity atlas for kenya helps to prioritise the emergency response on the most susceptible areas (see article, page 65). photograph: john tychsen, geus. 2. vibro-seismic data were acquired on stevns, denmark to be able to correlate the danian–campanian cores drilled along the east coast of stevns (see article, page 13). photograph: lars stemmerik, geus. 3. farming on the rooftops. experiment carried out at geus to investigate concentrations of pathogens leached to drainage water as a result of different slurry manure application methods. photograph: peter k. warna-moors, geus. 4. refuelling of helicopter on the arctic sea ice during seismic acquisition programme related to data collection for the continental shelf project around greenland. photograph: trine dahl-jensen, geus. frontispiece: facing page thrust sheet pair exposed in the rubjerg knude glaciotectonic complex, vendsyssel, denmark. photograph: stig a. schack pedersen, geus. chief editor of this series: adam a. garde scientific editors: martin sønderholm and a.k. higgins editorial secretaries: birgit eriksen and esben w. glendal external referees (numbers refer to first page of reviewed article): lars ole boldreel (29), michael houmark-nielsen (21, 61) and kristine thrane (25, 49, 53); geological institute, university of copenhagen, denmark. asger ken pedersen (37, 41, 45); geological museum, copenhagen, denmark. gregers dam (9, 13, 33) and michael larsen (17); dong energy, agern alle 24–26, hørsholm, denmark illustrations: : stefan sølberg with contributions from jette halskov lay-out and graphic production: annabeth andersen, carsten e. thuesen printers: schultz grafisk, albertslund, denmark manuscripts submitted: 15 january – 27 february 2006 final versions approved: 26 october 2006 printed: 29 november 2006 issn 1603-9769, 1604-8156 isbn-10: 87-7871-189-4 isbn-13: 978-87-7871-189-2 geological survey of denmark and greenland bulletin the series geological survey of denmark and greenland bulletin replaces geology of denmark survey bulletin and geology of greenland survey bulletin. citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 10, 68 pp. available from geological survey of denmark and greenland (geus) • øster voldgade 10 • dk-1350 copenhagen k • denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk or geografforlaget aps • rugårdsvej 55 • dk-5000 odense c • denmark phone: +45 63 44 16 83, fax: +45 63 44 16 97, e-mail: go@geografforlaget.dk © danmarks og grønlands geologiske undersøgelse (geus), 2006 4 7. review of survey activities 2005 a.a. garde 9. stratigraphy and palaeoceanography of upper maastrichtian chalks, southern danish central graben j.r. ineson, b. buchardt, s. lassen, j.a. rasmussen, p. schiøler, n.h. schovsbo, e. sheldon and f. surlyk 13. shallow core drilling of the upper cretaceous chalk at stevns klint, denmark l. stemmerik, f. surlyk, k. klitten, s.l. rasmussen and n. schovsbo 17. spit-systems – an overlooked target in hydrocarbon exploration: the holocene to recent skagen odde, denmark p.n. johannessen and l.h. nielsen 21. construction of 3d geological models in glacial deposits to characterise migration of pollution k.e.s. klint, f. von platen-hallermund and m. christophersen 25. advanced in situ geochronological and trace element microanalysis by laser ablation techniques d. frei, j.a. hollis, a. gerdes, d. harlov, c. karlsson, p. vasquez, g. franz, l. johansson and c. knudsen 29. east greenland and faroe–shetland sediment provenance and palaeogene sand dispersal systems m. larsen, c. knudsen, d. frei, m. frei, t. rasmussen and a.g. whitham denmark greece turkey albania ghana cameroon west indies uganda sudan tanzania mozambique latvia greenland madagascar canada united kingdom norway faroe islands kenya sweden rusland (novaya zemlya) geus working areas 2005. orange areas are covered in this volume. for further information on other working areas please refer to our website: www.geus.dk/international 5 33. continental crust in the davis strait: new evidence from seabed sampling f. dalhoff, l.m. larsen, j.r. ineson, s. stouge, j.a. bojesen-koefoed, s. lassen, a. kuijpers, j.a. rasmussen and h. nøhr-hansen 37. an integrative and quantitative assessment of the gold potential of the nuuk region, west greenland b. møller stensgaard, t.m. rasmussen and a. steenfelt 41. the tikiusaaq carbonatite: a new mesozoic intrusive complex in southern west greenland a. steenfelt, j.a. hollis and k. secher 45. archetypal kimberlite from the maniitsoq region, southern west greenland and analogy to south africa t.f.d. nielsen, m. jebens, s.m. jensen and k. secher 49. using zircon geochronology to resolve the archaean geology of southern west greenland j.a. hollis, d. frei, j.a.m. van gool, a.a. garde and m. persson 53. five slices through the nuussuaq basin, west greenland a.k. pedersen, l.m. larsen, g. krarup pedersen and k.s. dueholm 57. earthquake seismology in greenland – improved data with multiple applications t.b. larsen, t. dahl-jensen, p. voss, t. møller jørgensen, s. gregersen and h.p. rasmussen 61. radical past climatic changes in the arctic ocean and a geophysical signature of the lomonosov ridge north of greenland n. mikkelsen, n. nørgaard-pedersen, y. kristoffersen, s. juul lassen and e. sheldon 65. kensea – development of an environmental sensitivity atlas for coastal areas of kenya j. tychsen, o. geertz-hansen and j. kofoed india thailand lao pdr vietnam sri lanka mongolia 7 the present volume is the third issue of review of survey activities (rosa). it contains 15 four-page contributions that cover a wide range of the current activities at the geological survey of denmark and greenland (geus). thirteen of these are short scientific papers dealing with ongoing research by the survey and its external partners. for the first time the research-based papers in the review are now externally peer reviewed. a new standing panel of reviewers for rosa has been established to ensure that the contributions are of general interest to a wide readership and that, within the limitations of space, they maintain the normal scientific standards of the survey’s publications. all articles are planned to be easily readable by non-specialists, and since this is a review of survey activities, it should be borne in mind that many papers are first accounts of ongoing research. the fact that almost all contributions in the current volume are scientific in nature implies that while providing a timely panorama of current research at the survey, they are far from embracing all projects undertaken by the survey in denmark, greenland and other countries in 2005. a factual overview of the activities of geus as a whole can be obtained at geus’ website. in the present volume three papers deal with cretaceous–holocene onshore and offshore stratigraphy, sedimentology and palaeoceanography in denmark, in part related to hydrocarbon exploration. a fourth paper from denmark, that illustrates just one of the broad spectrum of the survey’s routine responsibilities undertaken on behalf of the state, addresses construction of 3d geological models to characterise the migration of point-source pollution in groundwater reservoirs. projects related to greenland and the arctic in general are represented in this volume by a group of nine papers. the first is a methodology paper describing advanced in situ geochronological and trace element microanalysis by laser ablation techniques, a now routine analytical tool at the survey that has provided data for several of the subsequent articles, including a study of sediment provenance in the east greenland – faroe islands – shetland region, and an account of zircon geochronology applied to archaean geological studies in southern west greenland. within the same region, a new method of integrative and quantitative assessment of the gold potential is presented, and two papers deal with newly discovered kimberlites and carbonatites and their potential economic significance. one paper describes five profiles through basalts and sedimentary rocks in the nuussuaq basin in west greenland, constructed using geological photogrammetrical techniques along coastal cliffs and steep valley sides. another paper presents new evidence for the presence of continental crust in the davis strait obtained from seabed sampling; this is an important new contribution to the long-standing debate of the nature of the crust under the labrador sea and davis strait and its stratigraphy. a report on ongoing studies of the deep crustal structure of greenland using earthquake seismology is presented, and a last paper concerning the arctic region describes radical former climatic changes in the arctic ocean and the geophysical signature of the lomonosov ridge north of greenland, and discusses the sensitivity of the sea-ice cover to global warming. the final paper in the present volume describes the development of an environmental sensitivity atlas for coastal areas of kenya. this project is just one of several current geus projects where the survey's broad technical and managing expertise is put to use in developing countries. review of survey activities 2005 adam a. garde chief editor © geus, 2006. geological survey of denmark and greenland bulletin 10, 7 only. available at: www.geus.dk/publications/bull geological survey of denmark and greenland bulletin 35, 2016, 75-78 75© 2016 geus. geological survey of denmark and greenland bulletin 35, 75–78. open access: www.geus.dk/publications/bull recent record-warm summers in greenland (khan et al. 2015) have started affecting the higher regions of the ice sheet (i.e. the accumulation area), where increased melt has altered the properties of firn (i.e. multi-year snow). at high altitudes, meltwater percolates in the porous snow and firn, where it refreezes. the result is mass conservation, as the refrozen meltwater is essentially stored (harper et al. 2012). however, in some regions increased meltwater refreezing in shallow firn has created thick ice layers. these ice layers act as a lid, and can inhibit meltwater percolation to greater depths, causing it to run off instead (machguth et al. 2016). meltwater at the surface also results in more absorbed sunlight, and hence increased melt in the accumulation area (charalampidis et al. 2015). these relatively poorly understood processes are important for ice-sheet mass-budget projections. regional climate models (rcms) simulate energy fluxes and mass transfer between the atmosphere and the icesheet surface. their accuracy depends on model physics and numerical sophistication, as well as on the atmospheric forcing implemented at their boundaries based on global weather reanalyses or general circulation models (see below). ice-sheet mass-budget calculations using rcms therefore need to be validated against observations. in this study, we evaluate the performance of the subsurface scheme of the hirham5 rcm (christensen et al. 2006) by comparing it with firn temperatures measured at the kan_u weather station from april 2009 to september 2013 (charalampidis et al. 2016). we determine the reasons for temperature biases by comparing hirham5 with a validated surface energy balance (seb) model over the same period (charalampidis et al. 2015). firn temperature measurements situated 1840 m above sea level (a.s.l.), kan_u is the uppermost automatic weather station at an elevation transect of meteorological and mass-budget monitoring sites in the south-western part of the greenland ice sheet (charalampidis et al. 2015; 67°0´n, 47°1́ w). the long-term equilibrium line altitude, where summer ablation balances winter accumulation, is 1553 m a.s.l. (van de wal et al. 2012). kan_u is located above that, and thus monitors melt, percolation and refreezing in firn. established in april 2009, the kan_u record includes the high melt seasons of 2010, 2011 and 2012 (charalampidis et al. 2015). the subsurface temperature analysis by charalampidis et al. (2016) revealed that in the 2010 and 2011 high melt summers, meltwater occupied the pore volume between 2 and 3 m below the surface (fig. 1a). the continued refreezing of this temporarily retained, near-surface liquid water until after the end of both the 2010 and 2011 melt seasons contributed to the merging of superimposed annual ice layers. these ice layers were observed at depths between 2.5 and 5.5 m in may 2012 (machguth et al. 2016). subsequently, meltwater by the end of august 2012 was confined in the upper 2.5 m relative to the may 2012 surface, with subsequent runoff in response to the intense surface lowering. by september 2012, after the onset of cold atmospheric conditions, refreezing occurred below 2.5 m by meltwater percolation to the limited available pore volume between the ice layers. the latent heat release by refreezing at depth in autumn 2012 resulted in a high december–january– february average firn temperature of –6.1°c between 2 and 5 m depth (charalampidis et al. 2016), while the accumulating snow cover provided thermal insulation from the cold winter atmosphere. the hirham5 regional climate model we use hirham5 at 5 × 5 km horizontal resolution, which has demonstrated good results for the climate of the greenland ice-sheet margin (e.g. langen et al. 2015). it uses 31 vertical atmospheric levels and a time step of 90 seconds. at the lateral boundaries, the model is forced at 6-hour intervals with wind, temperature, specific humidity and atmospheric pressure from the era-interim weather reanalysis (dee et al. 2011). the model computes processes in the atmosphere, including clouds, solar radiation attenuation, longwave radiation emission and precipitation. these variables then determine the energy balance and mass budget at the surface. daily-smoothed, modis-derived surface albedo regulates solar radiation absorption (box et al. 2012). regional climate-model performance in greenland firn derived from in situ observations charalampos charalampidis, dirk van as, peter l. langen, robert s. fausto, baptiste vandecrux and jason e. box 7676 the subsurface scheme (version 7.11) uses 25 layers with a total depth of 70 m water equivalent (c. 78 m physical distance at kan_u). it accounts for heat diffusion, vertical water transport and refreezing, as well as temperature, and pressure-dependent densification of snow and firn after vionnet et al. (2012). each layer can hold liquid water corresponding to 2% of the snow pore volume and excess water percolates downward to the next layer. water is assumed to run off when it encounters a layer of pore close-off density (i.e. 830 kg/m3; herron & langway 1980). before runoff occurs, the water is available for superimposed ice formation onto the ice layer. simulated versus observed firn temperatures the hirham5-simulated firn temperature evolution at kan_u is shown in fig. 1b. a seasonality following surface forcing is evident: 0°c from summer melting and about –20°c near the surface in winter. during the melt season, the simulation shows maximum firn temperature at depths c. 2 m in 2009 and c. 9 m in 2010, and even deeper in the following years. accordingly, the extent of the simulated temperate layer (i.e. temperatures between –1 and 0°c) increases from 6 m in summer 2009 to more than 10 m in 2010 and the following melt seasons. the propagation of the temperate conditions at depth suggests concurrent meltwater percolation, refreezing and latent heat release. however, the observed temperate layer did not extend below 3 m at any point (fig. 1a). the firn temperature bias (simulated minus observed) is shown in fig. 1c by comparing the interpolated hirham5 values at observational depths with the observed ones. the model bias is mostly positive and increases with depth. typical differences for the deepest measurements range between +6 and +12°c. negative differences occur during winter at depths less than 2 m in all years except 2012. in winter 2012, the rcm underestimates firn temperatures as deep as 3 m. table 1 shows the average summer and winter hirham5 firn temperatures at specific depths, and the biases. the summer difference averaged over all available depths is +5.7°c. better agreement between hirham5 and observations is found for winter with an average difference of +3.2°c. in winter 2012, hirham5 agreement is best, and is the only instance in the comparison when model bias at any of the listed depths was negative. this agreement is indicative of the abnormally warm conditions that persisted in the top 2–5 m firn after the extreme 2012 melt season, but also of the efficiency of the hirham5 simulation of surface-heat transfer into firn. explaining the bias subsurface differences between rcm and the observations can be due to differences in surface melt, quantity and depth of meltwater percolation and the timing of refreezing. b 2010 2011 2012 2013 si m u la te d d ep th f ro m s u rf ac e (m ) kan_u firn temperature 0 1 2 3 4 5 6 7 8 9 −20 −18 −16 −14 −12 −10 −8 −6 −4 −2 0 year 2010 2011 2012 2013 (°c) 2010 2011 2012 2013 (°c) (°c) a c si m u la te d d ep th f ro m s u rf ac e (m ) temperature bias 0 1 2 3 4 5 6 7 8 9 10 −20 −16 −12 −8 −4 0 4 8 12 16 20 d ep th f ro m s u rf ac e (m ) hirham5 firn temperature 0 1 2 3 4 5 6 7 8 9 10 −20 −18 −16 −14 −12 −10 −8 −6 −4 −2 0 fig. 1. a: firn temperatures observed at kan_u (charalampidis et al. 2016). b: hirham5-simulated firn temperatures at the location of kan_u, with the blue lines indicating the simulated depth of the mid-point of each layer. c: the difference between the two (b minus a). 77 the comparison of hirham5 with a validated seb model forced by in situ observation data (charalampidis et al. 2015) shows good agreement in simulated melt estimates (fig. 2a). hirham5 slightly underestimates melt in all years (differences less than 31 mj/m2) except 2011 (excess of 7 mj/m2). the cumulative difference in melt energy between the two models over the course of five melt seasons amounts to 68 mj/m2, approximately equal to the total melt in july 2013. this suggests that the positive firn temperature biases are not due to exaggerated melt. hirham5 substantially underestimates refreezing (fig. 2b). the differences are less than 15% in all years except 2012, when the difference is 44% and approximately equal to total refreezing in 2013 (380 kg/m2). with well-simulated melt and underestimated refreezing by hirham5, the firn temperature bias is due to prolonged wintertime refreezing at great depth. as a result of the overestimated latent heat release at depth during every year, hirham5 wintertime low temperature extremes remain at depths no greater than c. 3 m (fig. 1b). both hirham5 and charalampidis et al. (2015) overestimate the percolation depth. as analysed by charalampidis (2016), the seb model captures the thermal evolution of firn well. the model was initialised on 4 april 2009 based on firn temperature observations, and height-corrected 2012 firn densities, thus calculating realistic cold content (i.e. the required energy to raise firn temperature at 0°c) and heat diffusion estimates throughout the 4.5-year simulation. the seb model does not calculate liquid water retention, thus all percolating meltwater is refrozen at every time step, which is incorrect close to the surface (2–3 m depth; fig. 1a). however, no refreezing (i.e. no latent heat release) after the melt season at depth results in more realistic wintertime cooling of deep firn. hirham5 was initiated more than two decades before 2009. additionally, the current subsurface scheme of hirham5 cannot reproduce ice layers. this inability results in unrealistic representation of firn stratigraphy, and thus estimation of cold content, which is dependent on firn temperature and density. on 4 april 2009, the thermal state of the subsurface in hirham5 is the integrated result of all previous simulation years, and is already on average 4.6°c too warm in the upper 10 m of firn (fig. 1c). the associated cold content integrated over the first 10 m is 72 mj/m2. by summer temperatures (°c; june–july–august) 2 m –2.9 +3.1 –0.4 +4.3 –1.9 +3.5 –0.3 +5.3 –4.2 +3.5 3 m –3.6 +4.7 –1.0 +5.5 –2.3 +4.7 –1.1 +5.9 –4.4 +4.5 4 m –3.6 +6.1 –1.5 +6.5 –2.3 +6.0 –1.4 +6.9 –4.1 +5.1 5 m –3.6 +6.3 –1.7 +7.1 –2.1 +6.8 –1.3 – –3.6 +5.7 6 m –3.4 +6.7 –1.8 +7.9 –1.7 +7.7 –1.2 – –3.0 +6.4 following winter temperatures (°c; december–january–february) 2 m –9.5 +3.2 –8.9 +0.7 –12.2 +0.3 –8.7 –2.0 – – 3 m –6.5 +4.6 –5.9 +1.7 –8.4 +2.2 –6.0 +0.1 – – 4 m –4.6 +5.1 –3.8 +3.2 –5.5 +4.1 –4.0 +1.9 – – 5 m –3.4 +5.9 –2.3 +4.7 –3.3 +5.7 –2.6 +3.2 – – 6 m –2.7 +6.4 –1.2 +5.8 –1.9 +6.4 –1.6 – – – table 1. average hirham5 firn temperatures, linearly interpolated to specific depths (left columns) and biases (right columns) at kan_u (charalampidis et al. 2016) depth 2009 2010 2011 2012 2013 2010 2011 2012 2013 0 200 400 600 800 1000 year c u m u la ti ve m el t en er gy ( m j/ m 2 ) charalampidis et al. (2015) hirham5 a 2010 2011 2012 2013 0 c u m u la ti ve r ef re ez in g (1 0 3 k g/ m 2 ) charalampidis et al. (2015) hirham5 b 0.5 1 1.5 2 2.5 3 fig. 2. cumulative melt energy (a) and total refreezing (b) at kan_u by hirham5 and charalampidis et al. (2015). 7878 comparison, the cold content on the same day in charalampidis et al. (2015) is 151 mj/m2. one third of the difference in cold content is due to differences in firn density. near-complete cold content depletion (i.e. firn temperature at 0°c) of the first 10 m of firn is simulated by hirham5 for 2010. thereafter, this temperate firn retains liquid water, which refreezes during winter under the influence of subfreezing conditions diffused from the surface. this results in wintertime latent heat release at depth that in 2010 to 2012 is sustained until the beginning of the following melt season (fig. 2b). eventually, this premature depletion of cold content leads to overestimation of the percolation depth, liquid water retention and heat in firn. concluding remarks judging from the simulation of wintertime firn temperatures in the period april 2009 to september 2013, hirham5 is able to realistically reproduce subsurface processes when heat conduction dominates. yet the comparison of hirham5 with observations and seb model output reveals an overestimation of the percolation depth, liquid water retention and heat input from refreezing. for april 2009, hirham5 calculates less than half of the cold content estimate based on observations in the first 10 m of firn. this is the result of the 1989 initialisation of hirham5 and thus the cumulative effect of the imprecise determination of heat diffusion, as hirham5 is unable to adequately represent ice-layer formation in firn. a hirham5 subsurface scheme with improved accounting of shallow firn stratigraphy would greatly improve heat diffusion estimates over long simulation periods, and thus provide more reliable simulations. acknowledgements the kan_u weather station is funded by the greenland analogue project, with contributions from the programme for monitoring of the greenland ice sheet (promice). this is a promice publication and contribution number 76 of the nordic centre of excellence svali, ‘stability and variations of arctic land ice’, funded by the nordic toplevel research initiative (tri). the study has been supported by the danish council for independent research (dff) project 4002-00234 ‘understanding and predicting non-linear change in the permeability of greenland firn’. references box, j.e., fettweis, x., stroeve, j.c., tedesco, m., hall, d.k. & steffen, k. 2012: greenland ice sheet albedo feedback: thermodynamics and atmospheric drivers. the cryosphere 6, 821–839. charalampidis, c. 2016: climatology and firn processes in the lower accumulation area of the greenland ice sheet. digital comprehensive summaries of uppsala dissertations from the faculty of science and technology 1372, 81 pp. acta universitatis upsaliensis, uppsala, sweden. charalampidis, c., van as, d., box, j.e., van den broeke, m.r., colgan, w.t., doyle, s.h., hubbard, a.l., macferrin, m., machguth, h. & smeets, c.j. 2015: changing surface–atmosphere energy exchange and refreezing capacity of the lower accumulation area, west greenland. the cryosphere 9, 2163–2181. charalampidis, c., van as, d., colgan, w.t., fausto, r.s., macferrin, m. & machguth, h. 2016: thermal tracing of retained meltwater in the lower accumulation area of the southwestern greenland ice sheet. annals of glaciology, available on cjo2016, http://dx.doi. org/10.1017/aog.2016.2 christensen, o.b., drews, m., christensen, j.h., dethloff, k., ketelsen, k., hebestadt, i. & rinke, a. 2006: the hirham regional climate model version 5. danish meteorological institute technical report 06–17, 22 pp. copenhagen: danish meteorological institute. dee, d.p. et al. 2011: the era-interim reanalysis: configuration and performance of the data assimilation system. quarterly journal of the royal meteorological society 137, 553–597. harper, j., humphrey, n., pfeffer, w.t., brown, j. & fettweis, x. 2012: greenland ice-sheet contribution to sea-level rise buffered by meltwater storage in firn. nature 491, 240–243. herron, m.m. & langway, c.c. 1980: firn densification: an empirical model. journal of glaciology 25(93), 373–385. khan, s.a., aschwanden, a., bjørk, a.a., wahr, j., kjeldsen, k. & kjær, k. 2015: greenland ice sheet mass balance: a review. reports on progress in physics 78(4), 046801. langen, p.l. et al. 2015: quantifying energy and mass fluxes controlling godthåbsfjord freshwater input in a 5-km simulation (1991–2012). journal of climate 28, 3694–3713. machguth, h., macferrin, m., van as, d., box, j.e., charalampidis, c., colgan, w., fausto, r.s., meijer, h.a.j., mosley-thompson, e. & van de wal, r.s.w. 2016: greenland meltwater storage in firn limited by near-surface ice formation. nature climate change 6, 390–393. van de wal, r.s.w., boot, w., smeets, c.j.p.p., snellen, h., van den broecke, m.r. & oerlemans, j. 2012: twenty-one years of mass balance observations along the k-transect, west greenland. earth system science data 4, 31–35. vionnet, v., brun, e., morin, s., boone, a., faroux, s., le moigne, p., martin, e. & willemet, j.-m. 2012: the detailed snowpack scheme crocus and its implementation in surfex v7.2. geoscientific model development 5, 773–791. authors’ addresses c.c., d.v.a., r.s.f., b.v. & j.e.b., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k. e-mail: cc@geus.dk p.l.l., climate and arctic research, danish meteorological institute, lyngbyvej 100, dk-2100 copenhagen ø, denmark. c.c., also at: department of earth sciences, uppsala university, villavägen 16, se-752 36 uppsala, sweden. b.v., also at: arctic technology centre (artek), technical university of denmark, brovej, byg. 118, dk-2800 kgs. lyngby, denmark. geological survey of denmark and greenland. bulletin 10, 61-64 the arctic ocean is a landlocked basin, at present covered by perennial sea ice. during the past few decades a significant thinning and shrinking of the sea ice has been observed, and modelling studies indicate that the arctic ocean ice cover could, by the end of this century, almost disappear from most parts of the arctic ocean during peak summer seasons. it remains uncertain, however, whether the environmental changes are an enhanced greenhouse-warming signal or a result of natural (long-term) variability, but palaeoceanographic studies can contribute to our understanding of the natural variability of environmental parameters, e.g. sea-ice cover and oceanographic changes on time-scales of centuries to millennia. as part of the multidisciplinary eu project greenland arctic shelf ice and climate experiment (greenice), sediment coring and seismic reflection measurements have been undertaken in a hitherto unexplored part of the arctic ocean, the margin of the lomonosov ridge in the lincoln sea (fig. 1). the aim of the project was to study the structure and dynamics of the sea-ice cover and attempt to relate these to longer-term records of climate variability retrieved from sediment cores. the main field work was carried out in may 2004 from an ice camp established by a twin otter aircraft on drifting sea ice at 85°n, 65°w, c. 170 km north of alert, arctic canada. the camp was deployed over the shallowest part of the lomonosov ridge off the northern greenland/canada continental margin (fig. 1). the sea-ice drift would normally be between east and south, but persistent easterly winds resulted in a fast drift trajectory towards the wsw, such that the camp drifted a distance of approximately 62 km during the two weeks camp period. at present the study area is heavily ice covered, and forecast models of future shrinking arctic sea-ice cover suggest that this area is one of the least sensitive to warming in the arctic. the results obtained from the greenice project challenge this view. an unexplored area the reduction and thinning of arctic sea ice in recent decades (e.g. rothrock et al. 1999; acia 2004) has drawn attention to whether these environmental changes are an early reaction to global warming, or whether they are part of a long-term variation of the arctic environment. modelling studies of global warming effects indicate that the arctic is radical past climatic changes in the arctic ocean and a geophysical signature of the lomonosov ridge north of greenland naja mikkelsen, niels nørgaard-pedersen,yngve kristoffersen, susanne juul lassen and emma sheldon © geus, 2006. geological survey of denmark and greenland bulletin 10, 61–64. available at: www.geus.dk/publications/bull 500 600 700 800 9001000 1000 1100 12 00 #11 #10 76°w 74°w 72°w 70°w85°00´ 84°45´ 10 km arctic ocean lincoln sea alert lomonosov ridge fig. 1. upper: the greenice field camp area (marked by a red square) was deployed north of arctic canada and north greenland at the shallowest part of the submarine lomonosov ridge, in a region where no geologic record has hitherto been retrieved. lower: drift path of the field camp is shown by arrow and red line and coring stations by yellow and red dots. 61 likely to show a significant temperature increase, and that sea-ice cover could, by the end of this century, almost disappear during peak summer seasons (johannessen et al. 2004). such a scenario would not only have a dramatic impact on arctic ecosystems, navigation and indigenous people, but could also influence the thermohaline circulation and regional climate in the sub-arctic and north atlantic region. in a discussion of these scenarios, there is an urgent need for high-latitude arctic records of variations in climate, oceanography and sea-ice cover representing long time periods and, in particular, records of natural environmental change during earlier warm periods, which can be used to evaluate presentday changes. in spite of its importance, the recent geological record of many parts of the arctic ocean, including the lincoln sea, are still poorly known and hampered by difficult access. reduced ice cover during interglacial periods seismic data and sediment cores were collected from the drifting greenice station in this normally inaccessible area of the lincoln sea. during the camp period, 15 gravity core stations were established (figs 1, 2), and the retrieved cores were subsequently subjected to a wide array of investigations including ams-14c dating, faunal analysis of nannofossils and benthic and planktonic foraminifers, and stable isotope and geochemical analysis. the two longest cores, greenice core 10 (176 cm) and greenice core 11 (64 cm), show several characteristic colour cycles previously recorded in other parts of the arctic ocean (fig. 3; phillips & grantz 1997; nørgaard-pedersen et al. 1998; jakobsson et al. 2000; polyak et al. 2004; spielhagen et al. 2004). the stratigraphy of core 11 is based on nannoplankton, benthic foraminiferal assemblages and ams-14c dates and provides a record of the last c. 130 000 years, including the last interglacial period (eemian). preliminary investigations indicate that the longer core 10 contains a record of the last c. 200 000 years. planktonic foraminiferal assemblages are used as a key palaeoceanographic proxy, and a surprisingly large variability of these foraminifers was observed for an interior arctic ocean site. the discovery of abundant numbers of the small subpolar foraminifers turborotalita quinqueloba in two core sections, corresponding to the last interglacial and a younger warm interstadial (fig. 3), is an enigma, as this species indicates fairly strong subsurface atlantic water advection and possibly a much reduced summer sea-ice cover in the area compared to present-day conditions. the youngest part of the retrieved sediment record is condensed, but samples taken from close to the surface, representing holocene and recent conditions, lack the subpolar foraminifer species and thus indicate a consistent thick perennial sea-ice cover in accordance with present-day conditions (nørgaard-pedersen et al. in press) the results support the concept that interglacial conditions in the interior arctic ocean can vary considerably. at present, however, it is not known whether the influx of subpolar foraminifers was related to an ice-margin or polynyatype setting, or whether it reflects a generally reduced sea-ice cover of the interior arctic ocean. ongoing work aims to explore whether the observed trends can be traced to other key sites in the arctic ocean. seismic investigations and active faulting a 62 km long seismic reflection profile was collected during the drift of the greenice field camp (figs 1, 4). seismic reflection data were obtained from the shallowest part of the 62 a b fig. 2. a: lightweight gravity coring equipment (constructed by j. boserup, geus) used during the drift of the greenice camp. b: sediment cores were retrieved through a hole drilled in the ice. submarine lomonosov ridge facing the canadian/greenlandic continental margin, and comprise two parallel single channel lines (kristoffersen & mikkelsen 2006). the data reveal that the top of lomonosov ridge is bevelled at a water depth of 550 m and that only a thin sediment cover (less that 50 m) overlies the acoustic basement. pre-pleistocene sediments were probably eroded by a grounded marine ice sheet extending north from ellesmere island, and/or by deep draft icebergs. in the deep passage between the lomonosov ridge and the lincoln sea continental margin, more than 1 km of sediment is present. the uppermost 300 m of this succession reflects a significant sediment drift possibly related to 63 fig. 3. the greenice sediment cores show marked colour cycles. greenice core 11 covers a time span of c. 130 000 years and includes the eemian interglacial marine isotope stage 5e. abundant subpolar foraminifers (turborotalita quinqueloba) in eemian deposits indicate open water conditions not far from the greenice site. this is in contrast to holocene sediments that show a total dominance of polar species (neogloboquadrina pachydermal). 0 5 10 15 20 25 30 35 40 45 50 55 60 65 holocene polarí species subpolar’ species planktonic foraminifers nos/g sediment core 11 n. pachyderma (sin.) % t. quinqueloba nos/g sediment eemian interglacial 0 4000 8000 12000 16000 0 4000 80000 20 40 60 80 100 d ep th ( cm ) 1.0 1.5 2.0 2.5 3.0 2000 1500 1000 shot point 500 20 km t im e (s ec .) middle/upper pliocene – pleistocene upper cretaceous – neogene cretaceous upper jurassic – lower cretaceous basement sw ne lomonosov ridge fig. 4. seismic line retrieved during the ice drift over the lomonosow ridge north of canada and greenland (adapted from kristoffersen & mikkelsen 2006). 64 increased plio-pleistocene sediment input, and the underlying 700 m of sediment onlap a subsiding ridge slope. blocks of older margin sediments may represent the acoustic basement in the area. a basal unconformity, which may correspond to the hauterivian break-up unconformity of embry & dixon (1994), caps a series of nw–se-trending grabens, and several of the main graben faults extend to the sea bed and appear to have been active until recent times. acknowledgements the greenland arctic shelf ice and climate experiment (greenice) was supported by eu-grant evk2-2001-00280. references acia, 2004: impacts of a warming arctic – arctic climate impact assessment, 144 pp. cambridge: cambridge university press. embry, a.f. & dixon, j. 1994: the age of the amerasian basin. in: thurston, d.k. & fujita, k. (eds): 1992 proceedings, international conference on arctic margins, 289–294. anchorage, alaska, usa: u. s. department of the interior, minerals management service. jakobsson, m., løvlie, r., al-hanbali, h., arnold, e., backman, j.& mörth, m. 2000: manganese and color cycles in arctic ocean sediments constrain pleistocene chronology. geology 28, 23–26. johannessen, o.m. et al. 2004: arctic climate change: observed and modelled temperature and sea-ice variability. tellus a 56, 559–560. kristoffersen, y. & mikkelsen, n. 2006: on sediment deposition and nature of the plate boundary at the junction between the submarine lomonosov ridge, arctic ocean and the continental margin of arctic canada /north greenland. marine geology 225, 265–278. nørgaard-pedersen, n., spielhagen, r.f., thiede, j. & kassens, h. 1998: central arctic surface ocean environment during the past 80,000 years. paleoceanography 13, 193–204. nørgaard-pedersen. n., mikkelsen, n., lassen, s.j., kristoffersen, y. & sheldon, e. in press: arctic ocean sediment cores off northern greenland reveal reduced sea-ice concentrations during the last interglacial period. paleoceanography. phillips, r.l. & grantz, a. 1997: quaternary history of sea ice and paleoclimate in the amerasia basin, arctic ocean, as recorded in the cyclical strata of northwind ridge. geological society of america bulletin 109, 1101–1115. polyak, l., curry, w.b., darby, d.a., bischof, j. & cronin, t.m. 2004: contrasting glacial/interglacial regimes in the western arctic ocean as exemplified by a sedimentary record from the mendeleev ridge. palaeogeography, palaeoclimatology, palaeoecology 20, 73–93. rothrock, d.a., yu, y. & maykut, g.a. 1999: thinning of the arctic seaice cover. geophysical research letters 26, 3469–3472. spielhagen, r.f., baumann, k.-h., erlenkeuser, h., nowaczyk, n.r., nørgaard-pedersen, n., vogt, c. & weiel, d. 2004: arctic ocean deepsea record of northern eurasian ice sheet history. quaternary science reviews 23, 1455–1483. authors’ addresses n.m., n.n.-p., s.j.l. & e.s., geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350 copenhagen k, denmark; e-mail: nm@geus.dk y.k., department of earth science, university of bergen, allegaten 41, n-5007 bergen, norway. geological survey of denmark and greenland bulletin 7, 2004, p 1-7 geological survey of denmark and greenland bulletin 7 • 2005 review of survey activities 2004 edited by martin sønderholm and a.k. higgins geological survey of denmark and greenland danish ministry of the environment geological survey of denmark and greenland bulletin 7 keywords geological survey of denmark and greenland, survey organisations, current research, denmark, greenland. cover photographs from left to right 1. trenches in chalk at stevns, denmark, for acid leaching experiments simulating the reaction of hydrochloric acid in horizontal wells in north sea chalk reservoirs. photo: peter frykman. 2. as part of the greenice project, sea-ice thickness in the arctic ocean was measured using airborne helicopter-based electromagnetic profiling in 2004. the project is funded by the eu and carried out in co-operation between the geological survey of denmark and greenland (geus) and the alfred-wegener-institute for marine and polar research (awi) as well as other research institutes. photo: sibylle goebell, awi. 3. investigations related to geological storage of co2 from fossil fuels are becoming an increasingly important field of work for geus. photo: peter k. warna-moors. 4. geophysical investigations of the fault-related gold occurrences on the island of storø, not far from nuuk, the capital of greenland (see article on page 65). photo: peter w.u. appel. frontispiece: facing page the goldmine at nalunaq, south greenland. the mine opened in 2004 after more than ten years of intensive geological and technical investigations. photo: sven monrad jensen. chief editor of this series: adam a. garde scientific editors: martin sønderholm and a.k. higgins editorial secretaries: esben w. glendal and birgit eriksen illustrations: jette halskov lay-out and graphic production: annabeth andersen printers: schultz grafisk, albertslund, denmark manuscripts submitted: 7 december 2004 – 25 february 2005 final versions approved: 1 april 2005 printed: 29 july 2005 issn 1604-8156 issn 1603-9769 isbn 87-7871-164-9 geological survey of denmark and greenland bulletin the series geological survey of denmark and greenland bulletin replaces geology of denmark survey bulletin and geology of greenland survey bulletin. citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 7, 80 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk or geografforlaget aps rugårdsvej 55, dk-5000 odense c, denmark phone: +45 63 44 16 83, fax: +45 63 44 16 97, e-mail: go@geografforlaget.dk © danmarks og grønlands geologiske undersøgelse (geus), 2005 4 review of survey activities 2004 k. sørensen 7 oil generation from coal source rocks: the influence of depositional conditions and stratigraphic age h.i. petersen 9 identification of hydrocarbons in chalk reservoirs from surface seismic data: south arne field, north sea p. japsen, a. bruun, i.l. fabricius and g. mavko 13 forward modelling of seismic response from north sea chalk o.v. vejbæk and r. rasmussen 17 a revised lithostratigraphy for the paleogene – lower neogene of the danish north sea p. schiøler, j. andsbjerg, o.r. clausen, g. dam, k. dybkjær, l. hamberg, c. heilmann-clausen, l.e. kristensen, i. prince and j.a. rasmussen 21 seismology: neotectonics and structure of the baltic shield s. gregersen, m. glendrup, t.b. larsen, p. voss and h.p. rasmussen 25 new methods in provenance studies based on heavy minerals: an example from miocene sands in jylland, denmark c. knudsen, d. frei, t. rasmussen, e.s. rasmussen and r. mclimans 29 groundwater quality monitoring in denmark j. stockmarr 33 contents geus working areas 2004. orange areas are covered in this volume. 5 steam treatment of contaminated groundwater aquifers – development of pathogenic micro-organisms in soil c.s. jacobsen, s. elmholt, c.b. jensen, p.b. jakobsen and m. bender 37 field experimental design for pesticide leaching – a modified large-scale lysimeter b. nilsson, j. aamand, o.s. jacobsen and r.k. juhler 41 the storebælt gateway to the baltic j.b. jensen, o. bennike, w. lemke and a. kuijpers 45 quantifying the relationship between pollen sedimentation in lakes and land cover using historical maps a.b. nielsen 49 climatic warming: a trigger for glacial iceberg surges (‘heinrich events’) in the north atlantic? a. kuijpers, h. heinrich and m. moros 53 on-line presentation of mineral occurrences in greenland l. thorning, l.aa. christensen, b.m. nielsen, f. schjøth and h. stendal 57 precambrian mineralising events in central west greenland (66°–70°15´n) h. stendal and a.a. garde 61 greenstone belts in the central godthåbsfjord region, southern west greenland j.a. hollis, j.a.m. van gool, a. steenfelt and a.a. garde 65 detection of kimberlitic rocks in west greenland using airborne hyperspectral data: the hypergreen 2002 project t. tukiainen and l. thorning 69 prospecting for dimension stones in greenland t.v. rasmussen and h.k. olsen 73 small-scale mining – hazards and opportunities in kyrgyzstan and mongolia p.w.u. appel 77 7 in this second issue of review of survey activities the main themes of the survey’s work in denmark, greenland and elsewhere are presented for a professional, but not specialist readership. out of 18 articles, five report on greenland activities, and this reflects the obligation of the survey to spend approximately 30% of its finance law grant on greenlandrelated activities. the first four articles deal with petroleum-related matters and illustrate some of the pertinent questions concerning the prospective part of the north sea. although most oil generated in the north sea has its origin in marine shales of late jurassic age, oil and condensate have also been generated from paralic, coal-bearing sediments of middle jurassic age. most hydrocarbons in the danish north sea are found in chalk reservoirs, but paleocene reservoirs also contribute to danish oil production; both scenarios are covered by articles in this volume. as a result of the restructuring of governmental research organisations initiated by the present danish government, the seismological research – which until 2004 was part of the danish national survey and cadastre – has been moved to the geological survey of denmark and greenland (geus). the seismology group signals this move with a contribution to this year’s review concerning the baltic shield, of which eastern denmark forms a marginal part. geus has responsibility for monitoring groundwater quality throughout denmark. this responsibility carries with it an obligation for research into the processes affecting groundwater quality. this is an important field of research in a country with widespread and intensive agricultural activities, and that relies almost exclusively on groundwater for its water supply. a group of three reports in this volume witness to the range of important activities carried out by the departments of the survey working with ground water. quaternary research at the survey is presented in three articles that cover vegetation history, the evolution of the baltic following the last deglaciation and climate research in the north atlantic. the articles describing marine geology activities build on a long-standing co-operation between geus and the baltic sea research institute in warnemünde, germany. the survey’s activities in greenland in 2004 were concentrated in west greenland. three papers focus on investigations related to the mineral potential of the precambrian basement terranes, while one covers the possibility of exploiting dimension stones suitable for the international market. such studies cannot be carried out without a thorough regional knowledge of the geology of greenland. in 2004 geus reached a milestone in regional geological mapping of greenland when the last of 14 bedrock geology maps at 1:500 000, which together cover all of greenland’s ice free land areas, was published. the survey’s activities in denmark and greenland have now reached a similar stage, in the sense that in both countries resources spent on systematic mapping projects have now become subordinate to more focused activities. although the systematic regional mapping of greenland has reached its primary goal, this is not recorded by a separate article, and there are other activities in geus’ broad portfolio of tasks that, for one reason or another, are not covered by articles in this volume; to mention a few: greenland petroleum geology both onand offshore; surveys adjacent to the faroe islands and greenland directed towards the continental shelf project; ongoing mapping in greenland and denmark; projects relating to nationwide modelling of groundwater flow; work on geothermal resources and co2 sequestration. furthermore, the important activities dealing with data handling and the creation of new data in geus laboratories are not described as such. how these activities ultimately come together to the benefit of the general public and industry is, however, illustrated by the paper on on-line presentation of mineral occurrences in greenland. during 2004, geus carried out work in more than 20 countries outside denmark, greenland and the faroe islands (see map in table of contents in this issue). one such project related to small-scale mining in mongolia and kyrgyzstan is described. information on the survey’s activities not dealt with in this volume can be found on the geus website. review of survey activities 2004 kai sørensen director geological survey of denmark and greenland bulletin 20, 2010, 27–30 27 contamination of groundwater with pesticides and nitrate has compelled the danish government to launch a major hydrogeological mapping programme covering about 40% of the land area of denmark. numerous geophysical surveys are currently being carried out in order to acquire the necessary data. these new data are crucial for the 3-d geological models that are used in the planning of future water supply and landuse. normally, site-specific groundwater protection zones (thomsen et al. 2004) are based on groundwatermodelled catchment areas for each well, but proper 3-d geological models are needed in order to create a valid basis for the groundwater models. since most of the danish nearsurface geology is complex, a full geological understanding is required combined with in-depth interpretation of geological and geophysical data. much research has dealt with geophysical mapping and numerical groundwater modelling, but only limited research has combined these topics for geological modelling. prior to geophysical mapping, groundwater models were based on simple data extraction from well databases without inclusion of geophysical data. in the following, a concept for detailed 3-d geological modelling with hydrogeophysical data is presented for a specific area. data and study area the 150 km2 area for modelling is located in jylland (fig. 1). the surface is dominated by clayey tills and reaches an elevation of more than 160 m above sea level. prior to the geophysical survey, the only useful data for a geological model in this area came from boreholes. lithological descriptions were only available for 50–75% of the boreholes, and the boreholes were too widely spaced to provide an overview of the subsurface geology. hence a detailed geological model could not be constructed. two geophysical methods were employed in the model area, the airborne transient electromagnetic method (skytem; sørensen & auken 2004) and the pulled array continuous electrical sounding method (paces; sørensen 1996). the skytem survey was to map the deeper sections of the subsurface down to depths of 200 m, while the paces survey focused on the uppermost 25 m. both surveys were conducted along lines with a spacing of 250 m. the heterogeneity in the shallowest parts of the subsurface is commonly too great to obtain a proper correlation between the paces survey lines, but at larger depths a better correlation is achieved by the skytem data. by combining skytem and borehole data a solid basis for geological modelling was obtained in the study area. model concept a three-step approach for 3-d geological modelling used for hydrogeological purposes has recently been developed in denmark (jørgensen et al. 2008). with this approach the modelling both ensures a thorough data interpretation and utilises the potential that lies in establishing an understanding of the geological history. the approach divides the geological model into three submodels: 1. a general, conceptual geological model which is primarily descriptive and imaged in conceptual cross-sections. the conceptual geological model is based on a review of former work and existing literature. 3-d geological modelling of the egebjerg area, denmark, based on hydrogeophysical data flemming jørgensen, rasmus rønde møller, peter b.e. sandersen and lars nebel fig. 1. map of denmark showing the location of the study area, egebjerg, north of the city of horsens. 10°e 14°e 55°n 56°n 57°n sweden denmark germany 50 km jylland egebjerg study area horsens © geus, 2010. geological survey of denmark and greenland bulletin 20, 27–30. open access: www.geus.dk/publications/bull 2828 2. a 3-d geological model, which is intended to serve as a geological database for a given model area and may contain all kinds of geological information, both lithological and stratigraphical. the 3-d model is not necessarily established in areas with poor data coverage. 3. a 3-d hydrostratigraphic model, which is intended to serve as the input for numerical groundwater models. this model is based on the 3-d geological model but focuses on lithological and hydraulic parameters. normally it will cover the entire model space, incorporating a ‘best guess’ in areas with poor data coverage. traditional 3-d geological models for hydrogeological purposes are constructed as layer-cake models, where the layers are defined as the volume between two surfaces. elements in such models are thus defined by bounding surfaces. the surfaces are defined by digitised points and/or interpolated grids from the digitised points. such models do not require advanced modelling software, but critical restrictions may arise in areas of complex geology where layer variations may be difficult or impossible to model (turner 2006). as the geology in denmark is rarely organised as welldefined layers and hence cannot be properly described in a layer-based model, a different approach is needed. the challenge is to resolve heterogeneity to a degree that meets the demands of the user, e.g. numerical groundwater modelling. one way to do this is by ‘voxel’ (voxel = volumetric pixel) modelling (e.g. turner 2006). due to the high complexity in the egebjerg area voxel modelling has therefore been implemented and tested here. the voxel discretisation is 100 × 100 m in the x–y direction and 5 m in the z direction resulting in a 3-d voxel grid composed of 1.5 million voxels organised in 74 voxel layers. as the skytem data cover large parts of the area, a resistivity grid with discretisation and dimensions identical to the voxels is generated, basically supplying the voxel grid with a resistivity value (fig. 2). the voxel model is supplemented by a number of layer boundaries modelled as surfaces. these layer boundaries are based on digitised point swarms and interpolated to contoured 2-d grids. the software package geoscene 3d (i-gis 2010) is used for the modelling. the egebjerg model follows the three-step approach for geological modelling as described above, but minor adjustments were required in order to make use of the voxel model approach. the chosen model concept is as follows: 1. a general, conceptual geological model identical to the one described in the three-step approach above. 2. a 3-d stratigraphical model, in which the geological history of the model area is in focus. geological elements, structures and stratigraphical boundaries are modelled and subdivided according to their origin. the geological elements comprise boundaries such as top chalk, top palaeogene, erosional unconformities, stratigraphical units such as palaeogene and miocene, buried tunnel valleys, glaciotectonic complexes and interglacial units. this model is a combined layer-based and voxel model and is generally comparable to the 3-d geological model within the three-step approach. fig. 3. stratigraphical log for the model area. the thickness of the entire section corresponds to about 200 m. for explanation see the text. upper glacial sequence interglacial lake sediments lower glacial sequence miocene palaeogene danian fig. 2. selected section of the skytem resistivity voxel grid. cell size: 100 × 100 × 5 m. resistivity 100 ohmm1 10 29 3. a 3-d lithofacies model that is entirely voxel based, and in which all voxels are allocated an attribute for lithofacies. the model is based on the 3-d stratigraphical model. a challenge here is to estimate the lithofacies in heterogeneous areas or areas with poor data coverage. a qualitative approach for uncertainty assessment of the interpretations is used and attributed to each voxel in order to visualise the overall model uncertainties (sandersen 2008). this is important in connection with the subsequent implementation in a numerical groundwater model. the uncertainty approach is also applied to the 3-d stratigraphical model. model results the geological modelling resulted in the stratigraphical log shown in fig. 3. it comprises a tertiary sequence followed by glacial and interglacial sediments above the pre-quaternary surface. this surface unconformably cuts layers of miocene clay, silt and sand, and palaeogene clay, which in turn overlie danian limestone at depths of about 200 m. only one borehole in the area reaches the limestone, but in some areas, where the skytem soundings penetrate the palaeogene clay, the limestone is seen as a layer with slightly elevated resistivity values. the palaeogene clay shows high electrical conductivity and responds as a distinct and well-defined surface in the skytem data. erosional remnants of miocene deposits are only found in minor parts of the study area. a model section showing parts of the pre-quaternary sequence is shown in fig. 4. the quaternary sequence is divided into an upper glacial/ interglacial sequence and a lower glacial sequence. several buried valleys are present within the lower glacial sequence, and at least four generations occur here. the valleys are interpreted as tunnel valleys (sensu jørgensen & sandersen 2006), and the sequence comprises a complex setting of cross-cutting valleys that have repeatedly been incised and filled. the valleys are primarily filled with till and glaciolacustrine clay but coarse meltwater deposits also occur. the valleys are modelled by voxels due to their high spatial complexity (fig. 5). in the skytem data, the lower glacial and upper glacial/ interglacial sequences are divided by an apparently widespread, gently undulating resistivity boundary that is considered to be an erosional unconformity. this unconformity appears as a widespread and more or less horizontal contrast in the resistivity data; with alternating high and low resistivity values on each side of the boundary. above the unconformity a unit of interglacial diatomaceous lake sediments is found in boreholes. pollen analyses show that these sediments were deposited during the holsteinian interglacial (odgaard 2010), and the erosional unconformity and the glacial sequence below, including most of the tunnel valleys, are therefore of elsterian age or older. some subareas in the study area are glaciotectonically deformed. this is seen where tertiary clay is found above glacial sediments in the boreholes. these large rafts of palaeogene clay, only indicated by a few boreholes, are well resolved in the skytem data. the deformed layers are in some places found inside the buried tunnel valleys showing that the deformation took place after the valley formation. the 3-d stratigraphical model as described above is used as the basis for the construction of the 3-d lithofacies model. some units in the 3-d stratigraphical model are more or less directly converted to lithofacies, but other units are subdivided, merged or reordered prior to incorporation into the 1 4 3 2 fig. 4. 3-d stratigraphical model section showing from below the boundaries of top chalk (1: green surface), top palaeogene (2: lower grey surface), an internal miocene boundary (3: middle grey surface) and the prequaternary surface (4: upper grey surface). boreholes and the skytem resistivity grid are also shown. vertical exaggeration: 5 times. fig. 5. a series of buried tunnel valleys modelled by voxels in the 3-d stratigraphical model: each colour represents one buried tunnel valley. the palaeogene surface is shaded in grey. vertical exaggeration: 5 times. 3030 3-d lithofacies model. a selected part of the model and its corresponding uncertainty is shown in fig. 6. concluding remarks in the egebjerg study area detailed stratigraphical and lithological information has been obtained by studying 3-d resistivity grids based on the skytem data and evaluating these against borehole data. in order to make full use of the new information for numerical groundwater modelling or for other purposes, the data have been incorporated in a 3-d geological model. the 3-d modelling was carried out in two steps, where two differently focused 3-d models were constructed. first a 3-d model focusing on stratigraphy was constructed and subsequently, based on this, a 3-d lithofacies model was constructed. this approach enables in-depth geological interpretation and secures maximum utilisation of the large data sets. the heterogeneous geology revealed by the skytem data cannot be sufficiently incorporated into a simple layer-based model, and a combination of a voxel and a layer model has therefore proven successful for the egebjerg study area. acknowledgements the project is supported by environment centre aarhus. we thank klaus petersen and stine rasmussen for helpful discussions. references i-gis 2010: geoscene 3d. http://www.i-gis.dk/default.aspx?tabid=132. jørgensen, f. & sandersen, p.b.e. 2006: buried and open tunnel valleys in denmark – erosion beneath multiple ice sheets. quaternary science reviews 25, 1339–1363. jørgensen, f., kristensen, m., højberg a.l., klint, k.e.s., hansen, c., jordt, b.e. richardt, n. & sandersen, p. 2008: opstilling af geologiske modeller til grundvandsmodellering. geo-vejledning 3, 176 pp. copenhagen: geological survey of denmark and greenland. odgaard, b. 2010: pollenanalytisk datering af ferskvandsaflejring i dgu nr. 107.733, horsens vandværk, 2 pp. unpublished report, aarhus universitet, danmark. sandersen, p.b.e. 2008: uncertainty assessment of geological models – a qualitative approach. in: refsgaard, j.c. et al. (eds): calibration and reliability in groundwater modelling: credibility of modelling. international association of hydrological sciences publication 320, 345–349. sørensen, k. 1996: pulled array continuous electrical profiling. first break 14, 85–90. sørensen, k.i. & auken, e. 2004: skytem – a new high-resolution helicopter transient electromagnetic system. exploration geophysics 35, 191–199. thomsen, r., søndergaard, v.h. & sørensen, k.i. 2004: hydrogeological mapping as a basis for establishing site-specific groundwater protection zones in denmark. hydrogeology journal 12, 550–562. turner, a.k. 2006: challenges and trends for geological modelling and visualisation. bulletin of engineering geology and the environment 65, 109–127. authors’ addresses f.j. & r.r.m., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: flj@geus.dk p.b.e.s., grontmij | carl bro a/s, dusager 12, dk-8200 århus n, denmark. l.n., i-gis, voldbjergvej 14, 2., dk-8240 risskov, denmark. a b fig. 6. selected section of the 3-d lithofacies model (identical to fig. 2). a: the lithofacies voxel model. blue: palaeogene clay. dark brown: clay till. light brown: sand till. red: meltwater sand/gravel. light red: sand. grey: diatomite. b: uncertainty of the lithofacies model. light grey: low uncertainty. grey: medium uncertainty. dark grey: high uncertainty. vertical exaggeration: 3 times. geological survey of denmark and greenland bulletin 35, 2016, 51-54 51© 2016 geus. geological survey of denmark and greenland bulletin 35, 51–54. open access: www.geus.dk/publications/bull the chemical composition of chalk and marl reflects the mixture of carbonate particles and clastic input deposited on the seabed together with growth of authigenic minerals and diagenesis. the rørdal quarry in jylland (fig. 1) is known for its alternating chalk–marl succession (surlyk et al. 2010) and the aim of this article is to investigate how this cyclicity is reflected in the geochemical signature of the sequence and test if this has implications for the interpretation of the depositional environment as well as the chemostratigraphy in the chalk. the observed variation in the benthic fauna and the cyclic character of the chalk–marl succession may reflect an environmental response to orbital forcing. the benthic fauna shows higher species diversity and density in the chalk than the marl layers, suggesting more favourable living conditions in the former (lauridsen & surlyk 2008), whereas no differences in environmental stress between the two environments could be derived from a study of the trace fossils (lauridsen et al. 2011). the c. 10 m thick rørdal member was established as a lithostratigraphic unit by surlyk et al. (2010) and recognised as an expression of a late maastrichtian cooling event. x-ray diffraction (xrd) analysis indicates that the marly layers represent a relative increase in smectite clays and illite together with some quartz and analcime. the inorganic geochemical evolution in maastrichtian chalk was investigated by jørgensen (1986) who noted changes in sr/ca and mn/ca ratios towards the end of the maastrichtian in the north sea central graben, leading to the conclusion that geochemistry is a “conceivable tool in basin analysis in the lithologically rather monotonous chalk sequence” (p. 267). however, a systematic geochemical analysis of the chalk has never been undertaken and this work is an attempt to look into the potential of understanding the geology of the chalk by analysing a wide range of elements in chalk and marl. chemical variability related to clay content sixty-three samples from the rørdal quarry were ground, digested in aqua regia and analysed for 66 elements using the elan 6100 quadrupole icp-ms at the geological survey of denmark and greenland (geus). the data reduction was based on totalquant software with emphasis on the trace element analysis. the analytical procedure follows larsen et al. (2009). the cyclicity and the alternation between chalk and marl are easily recognised in the geochemical profile (fig. 2) with an increase in the concentration of major and minor elements such as al, si, fe, mg and k in the marly layers caused by the presence of clay. the eight peaks in fig. 2 can be correlated to the eight peaks in the marly layers of the rørdal member which can be identified in the gamma log in a nearby well (surlyk et al. 2010). the gamma-radiation in the well log can be explained by the increased content of k in the marl (fig. 2a). iron is the element after al with the highest increase in concentration in the marly layers indicating that the clay minerals are rich in fe. magnesium also shows an increase in the marly layers (fig. 2a), and the combined clay minerals in the marl – which according to surlyk et al. (2010) consist of smectite and illite – must be rich in fe, k and mg. the content of al, fe and k is about three times higher in the marl than in the chalk. geochemistry of the maastrichtian rørdal member, jylland, denmark: ce anomaly as a palaeo-redox proxy christian knudsen and bodil w. lauridsen denmark sweden 50 kmgermany jylland møns klint stevns klint rørdal quarry n lake profile 1 km 10°e 55° 57°n a l o w er u p p er m aa st ri ct h ia n p al eo ce n e d an ia n fiskeler mb rørdal mb c b fig. 1. a: location of the rørdal quarry, stevns klint and møns klint in denmark. b: map of the rørdal quarry with the profile where the samples were collected. c: stratigraphic section with rørdal and fiskeler members. 5252 the concentration of many trace elements, including the rare-earth elements (ree), is also elevated in the marly layers (fig. 2b). if all ree were located in clay the increase would be three times higher, as is the case with al, k and fe, but the content of ree in the marl is ‘only’ 1.5 times higher than in the chalk. this suggests that although the ree concentration is higher in the clay than in the chalk a substantial fraction of the rees is located in the carbonate component of both chalk and marl. lanthanum and nd follow parallel tracks (fig. 2b), whereas ce crosses both tracks with relatively high ce in the marl. light ree (la, ce, pr, nd and sm) are more abundant than heavy ree (er, tm, yb and lu) in the marl compared to the chalk, which can also be seen as higher la/lu ratios (figs 3e, 4). lithophile elements such as th, cs, rb, be, cr, sc and li increase in concentrations in the marl (of which rb and li are shown in figure 2c) and are c. 2.5 to 3 times higher in the marl than in the chalk, similar to the behaviour of al and k. this suggests that these elements are almost entirely located in the clay. chalcophile elements such as pb, as, zn and cu (fig. 2d) are c. 1.5 times higher in the marl than in the chalk similar to the ree mentioned above. the content of most elements in the succession varies with the clay content. however, this is not the case for nb, zr, mn and ti (fig. 3a). mica and clay commonly contain some ti, zr and nb, elements that also occur in mafic volcanic material. these elements are difficult to leach out of minerals and generally follow clay and mica in sedimentary environments. if the clay was clastic or volcanogenic one would expect that the content of elements such as ti, zr and nb would be elevated in the marl. however, this is not the case, suggesting that a fraction of the clay may be of authigenic origin. manganese is probably hosted in the carbonate and its concentration is relatively stable throughout the analysed section. there is a decrease in the ba content in the rørdal member (fig. 3b) relative to the underlying hvidskud member with the lowest contents in its middle part. there is an increase in sr content in the overlying sigerslev member. jørgensen (1986) observed a similar increase in the sr content in chalk from the north sea basin and this may be regional. ce anomaly as a proxy for redox potential the ree commonly occur as cations with a valency of three and the different ree have similar geochemical behaviour in the sedimentary environment. however, one of the elements, ce, can also occur as ce4+ with a very low solubility in marine environments as compared to ce3+. where no fractionation of ce relative to the other ree has occurred, the distribution would be seen as a straight line on fig. 3d. to quantify the fractionation of ce, the term ce anomaly (ce/ce*) has been introduced; ce is the measured cerium concentration and ce* is what the ce content would have been without fractionation, based on the contents of la and nd. the ce anomaly is calculated: ce/ce* = 3cen/(2lan + ndn) where reen is the measured concentration normalised relative to chondrite (boynton 1984). in this paper a ce anomaly is referred to as negative when ce is depleted relative to the other ree and as positive if ce is enriched relative to the other ree. in 0 1 0 0 0 2 0 0 0 3 0 0 0 4 0 0 0 5 0 0 0 k m g fe a l si a ppm 1 .0 1 0 b y b d y g d sm n d p r c e l a 0 .1 1 .0 1 0 c r b g a c r vsc l i 0 .1 1 0 1 .0 d p b a s c u n i c o z n 0 5 10 15 m ( h ei gh t) u p p er m aa st ri ch ti an r ø rd al m b si ge rs le v m b h vi d sk u d m b fig. 2. chemical profiles through the rørdal member including the uppermost part of the hvidskud member and the lowermost part of the sigerslev member. marly layers are shown as grey shaded bands. a: major and minor elements. b: ree; logarithmic scale. c: lithophile trace elements; logarithmic scale. d: chalcophile elements with ni and co; logarithmic scale. stratigraphy from surlyk et al. (2013). 53 the modern marine environment, a negative ce anomaly is indicative of oxic conditions (german & elderfield 1990) which, in turn, is caused by the low solubility of ce4+ and low availability of ce. this feature offers the possibility to look into variations in the redox conditions during deposition of the chalk (jeans et al. 2015). in fig. 4 it can be seen that there is a negative ce anomaly both in the chalk and in the marl in the rørdal quarry as well as in the chalk and the fish-clay at stevns. there is also a negative eu anomaly which is probably caused by the source of the ree being slightly depleted in eu (frei & frei 2002). the negative ce anomaly in the chalk at stevns is more pronounced than in the chalk at rørdal (fig. 4) suggesting that the environment during deposition of the chalk at stevns was more oxic than at rørdal. the chalk at stevns has a higher stratigraphic position (in the sigerslev member) and the change in the ce anomaly with time could suggest that the oxygen availability in the chalk sea increased with time during the late maastrichtian. alternatively, this could suggest that the chalk at stevns was deposited in a shallower sea. the ce concentration at rørdal changes with lithology and stratigraphic position in the section (fig. 3d); it is lower in the chalk than in the marl, indicating that the environment was more oxic in the chalk than in the marl. the ree (including ce) are located in the carbonate component as well as in the clay as mentioned above. accordingly, the ce anomaly reflects the combined effect of the composition of the seawater where the coccoliths and where the clay were formed. surlyk et al. (2010) suggested that the clay in the rørdal member is derived from volcanic eruptions. such clay, derived from alteration of volcanic glass, is likely to have an overall resemblance to the ree distribution in the source. however, the transformation from glass to clay in the marine environment could be the cause of the negative ce anomaly found in the marl. the fish-clay at stevns also has this negative ce anomaly which is indicative of formation of clay in a marine environment (kastner et al. 1984; frei & frei 2002). the question is then whether the clay was (trans)formed from volcanic glass in the free water masses or in the seabed. the seabed is likely to have been less oxic than the free water mass, and formation of the clay in the seabed would explain the difference in the redox potential compared to the chalk formed in the free water mass. finally, the composition of the clay minerals could have been modfied during diagenesis in the presence of organic material affecting the ce anomaly. 0 .0 1 1 .0 1 0 0 a n b z r r b m n t i ppm 0 2 0 4 0 b b a 0 5 0 0 1 0 0 0 1 5 0 0 c sr 0 0 .2 0 .4 0 .6 d ce anomaly 4 0 8 0 1 2 0 e l a/ l u 0 10 15 m ( h ei gh t) 5 u p p er m aa st ri ch ti an r ø rd al m b si ge rs le v m b h vi d sk u d m b fig. 3. chemical profiles in the rørdal member. same stratigraphy as fig. 1. marly layers are shown as grey shaded bands. a: nb, zr, rb, mn and ti; logarithmic scale. b: ba. c: sr. d: ce anomaly. e: la/lu ratio; logarithmic scale. fish clay, stevns chalk, stevns marl, rørdal chalk, rørdal la ce nd sm gd dy er yb lupr pm eu tb ho tm 1 0 sa m p le /r e e c h o n d ri te 1 0 0 fig. 4. chondrite-normalised ree distribution patterns for the rørdal samples compared with ree data from stevns klint (frei & frei 2002). 5454 summary and outlook the chemical composition of the chalk–marl sequence at rørdal in jylland reflects the varying proportions of carbonate and clay. the content of major and minor elements such as si, al, fe and k is proportional to the clay content, with a threefold increase in the marly layers relative to the chalk layers. it is suggested that these elements are located in clay minerals together with lithophile trace elements such as li, ga, rb, cs and th which likewise increase threefold in the clay relative to the chalk. other elements such as pb, as, zn and cu and the ree are found in c. 1.5 times higher concentrations in the marl, and these elements are located both in the clay and the carbonate component of the succession. the clay has a light ree enrichment compared to the chalk. formation of the clay as authigenic minerals may explain why large ion lithophile elements such as zr and nb, as well as ti, are not incorporated in the clay. the content of ba in the rørdal member is low relative to the underand overlying strata, and the deposition of ba could be tested as a chemo-stratigraphic marker and environmental proxy. the ree distribution shows a negative ce anomaly both in the chalk and in the marl, suggesting that the environment was oxic throughout the deposition, and as the ce anomaly is more pronounced in the chalk, the environment was more oxic where the carbonate was formed than in the clay. this observation matches the observation that there is higher species diversity and density in the chalk than in the marl (lauridsen & surlyk 2008). if the clay and the carbonate were formed in the same water mass, the observed alternating redox level would be indicative of changing conditions in the sea. however, if the main part of the clay was formed in the seabed as authigenic minerals or was affected by diagenesis, then the alternating redox levels indicated by the ce anomaly reflects changes in redox conditions at or in the seabed. the ce anomaly in chalk from stevns is larger than in chalk from rørdal, suggesting that the environment was more oxic in the chalk sea towards the end of the maastrichtian. these differences could be related to changes in temperature over time. a recent study by thibault et al. (2016) suggests a global cooling with superimposed cool/warm fluctuations in the last 8 ma of the maastrichtian including e.g. the rørdal and sigerslev members. it is suggested that the ce anomaly can be used as a palaeo-redox indicator or proxy to compare environments across chalk basins. in this context it is interesting to analyse both carbonate and clay components from the marl. acknowledgement we wish to thank lars stemmerik for constructive comments on earlier versions of the manuscript. references boynton, w.v. 1984: cosmochemistry of the rare earth elements: meteorite studies. in: henderson, p. (ed.): rare earth element geochemistry, 63–114. amsterdam: elsevier. frei, r. & frei, k.m. 2002: a multi-isotopic and trace element investigation of the cretaceous–tertiary boundary layer at stevens klint, denmark – inferences for the origin and nature of siderophile and lithophile element geochemical anomalies. earth and planetary science letters 203, 691–708. german, c.r & elderfield , h. 1990: application of the ce anomaly as a paleoredox indicator: the ground rules. paleoceanography 5, 823–833. jeans, c.v., wray, d.s. & williams, c.t. 2015: redox conditions in the late cretaceous chalk sea: the possible use of cerium anomalies as palaeoredox indicators in the cenomanian and turonian chalk of england. acta geologica polonica 65, 345–366. jørgensen, n.o. 1986: geochemistry, diagenesis and nannofacies of chalk in the north sea central graben. sedimentary geology 48, 267–294. kastner, m., asaro, f., michel, h.v., alvarez, w. & alvarez, l.w. 1984: the precursor of the cretaceous–tertiary boundary clays at stevns klint, denmark, and dsdp hole 465a. science 226 (4671), 137–143. larsen, l m., heaman, l.m., creaser, r.a., duncan, r.a., frei, r. & hutchinson, m. 2009: tectonomagmatic events during stretching and basin formation in the labrador sea and the davis strait: evidence from age and composition of mesozoic to palaeogene dyke swarms in west greenland, journal of the geologicla society (london) 166, 999–1012. lauridsen, b.w. & surlyk, f. 2008: benthic faunal response to late maastrichtian chalk–marl cyclicity at rørdal, denmark. palaeogeography, palaeoclimatology, palaeoecology 269, 38–53. lauridsen, b.w., surlyk, f. & bromley, r.g. 2011: trace fossils of a cyclic chalk–marl succession; the upper maastrichtian rørdal member, denmark. cretaceous research 32, 194–202. surlyk, f., stemmerik, l., ahlborn, m., harlou, r., lauridsen, b.w., rasmussen, s.l., schovsbo, n., sheldon, e. & thibault, n.r. 2010: the cyclic rørdal member – a new lithostratigraphic unit of chronostratigraphic and palaeoclimatic importance in the upper maastrichtian of denmark. bulletin of the geological society of denmark 58, 89–98. surlyk, f., rasmussen, s.l., boussaha, m., schiøler, p., schovsbo, n.h., sheldon, e., stemmerik, l. & thibault, n.r. 2013: upper campanian–maastrichtian holostratigraphy of the eastern danish basin. cretaceous research 46, 232–256. thibault, n., harlou, r., schovsbo, n.h., stemmerik, l., surlyk, f. 2016: late cretaceous (late campanian–maastrichtian) sea surface temperature record of the boreal chalk sea. climate of the past 12, 429–438. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark, e-mail: ckn@geus.dk geological survey of denmark and greenland bulletin 35, 2016, 47-50 47© 2016 geus. geological survey of denmark and greenland bulletin 35, 47–50. open access: www.geus.dk/publications/bull marine interglacial deposits are fairly common and widespread in denmark, but so far none have been reported from the herning area in central jylland. in 2014, the geological survey of denmark and greenland (geus) received samples at one metre intervals from a borehole at 55°59.3´n, 8°56.6́ e (elevation 27.56 m above sea level), at hesselvigvej 7 near kibæk in central jylland (fig. 1). the succession consisted of miocene and quaternary deposits. the quaternary part was dominated by glaciofluvial sand and a single till bed, but it also contained a marine clay unit (16–21 m depth, 6.5–11.5 m a.s.l.). this marine clay contained spines of the sea-urchin echinocardium cordatum, a boreal species known from eemian and holocene deposits from denmark, but unknown from interstadial deposits. a lacustrine unit between 26 and 33 m depth (5.5 m b.s.l. to 1.5 m a.s.l.) consisted of clay, calcareous-rich gyttja and diatomite. because quaternary marine deposits are unexpected in this part of denmark, we report here on analyses of pollen from the lacustrine unit and foraminifera from the marine unit, and we compare these with some interglacial records from jylland. the analysed samples were treated with standard laboratory methods. middle pleistocene interglacial deposits near herning, jylland, denmark bent v. odgaard, karen l. knudsen, ole bennike and henrik j. granat fig. 1. a: map of denmark showing the location of the herning area (red rectangle) and selected interglacial sites. 1: harboøre (dgu 43.75), 2: kås hoved (dgu 45.759), 3: holstebro nord (dgu 64.248), 4: hoven (dgu 103.1011), 5: vorbasse (dgu 123.1217). b: map of the herning area showing interglacial sites. c: simplified lithological log of the core at kibæk (dgu 95.2875), with two interglacial units. b lind 10°e 100 km sweden jylland germany 57°n 55°n a dgu 95.2875 (kibæk) 1 2 3 4 5 b interglacial lake deposits interglacial marine clay glaciofluvial sand glaciofluvial sand glaciofluvial sand glaciofluvial sand miocene sand miocene clay clayey till c d ep th b el o w t er ra in s u rf ac e (m ) 0 10 20 30 40 50 60 70 herning harreskov 9°e 56°n 5 km core open sections denmark 4848 the lacustrine unit pollen data from two samples (at 31 and 30 m depth) from the organic lacustrine (gyttja) unit clearly show that this is an interglacial deposit with thermophilous taxa such as tilia (lime) and buxus (box). the lack of carpinus (hornbeam) is noteworthy and makes a correlation with the holstenian or eemian periods unlikely. in 2003 a core was secured through a thick interglacial deposit at nearby lind, 14 km north of the kibæk borehole (kronborg & odgaard 2004; fig. 1). the pollen record of the interglacial part of the lind core shows a clear correlation to the harreskovian (andersen 1965), with absence of carpinus, presence of picea (spruce) throughout the series and traces of celtis (hackberry; fig. 2). an ordination (principal component analysis, pca) on the lind interglacial pollen samples with the kibæk pollen spectra as supplementary samples shows that the bottom kibæk sample correlates well with a level of 28.4 m of the lind core and the upper kibæk sample with lind at 23.4 m (figs 2, 3). this indicates that a full, but thin interglacial lacustrine succession is present at kibæk. the reason for the thin succession could be that the core site is located near the margin of the former lake basin. the marine unit the foraminiferal assemblages from the marine unit (fig. 4) are dominated by elphidium excavatum, with ammonia beccarii second in abundance. the species composition indicates subtidal, inner-shelf conditions with a gradual shallowing of the water depth and deposition during a full interglacial period. the species ammonia beccarii currently has a northern geographical range limit along southern norway, and it is not present in interstadial deposits such as the bølling and allerød in denmark. ammonia beccarii immigrated with the marine transgression in the early holocene. a principal component analysis (pca) on selected interglacial foraminiferal stratigraphies from central and northfig. 2. simplified percentage pollen diagram from the interglacial deposit at lind with indication of the stratigraphical correlation of the two pollen samples from kibæk (dgu 95.2875; purple lines). fig. 3. principal component analysis (pca) biplot of the interglacial pollen sequence of lind with the two kibæk samples as supplementary samples not influencing the geometry of the plot. the percentage values were log-ratio transformed prior to the analysis. pca is an ordination technique, which reduces a multidimensional space (here of pollen types) with associated frequencies to fewer dimensions (pca axes) in such a way that the new dimensions are determined by the directions of largest variation in the original data set. the technique also allows easy comparison between samples of multivariate data such as those derived from microfossil analysis. 23 24 25 26 27 28 29 30 0 200 20 40 0 20 40 0 20 0 20 0 20 0 20 0 20 0 0 0 0 20 0 20 percentages pin us d ep th ( m ) bet ula sal ix pic ea aln us quer cus co rylu s ulmus to ta l tili a c ord ata ty pe tax us ce ltis dryo pte ris ty pe po ac ea e fraxi alnus quercus tilia ulmus corylus taxus bruckenthalia calluna fraxinus poaceae filipendula pinus picea betula artemisia salix –1.1 kibæk samples species cyperaceae lind – 1 .0 1 .0 1.1 49 western jylland shows the kibæk samples to be intermediate between two long records from kås hoved and holstebro (fig. 5). this position probably reflects an intermediate facies of the kibæk deposit between the relatively shallow kås hoved facies with haynesina orbiculare as an important species, and the more open marine environment at holstebro with elphidium margaritaceum as one of the indicators of relatively high salinity. the kibæk foraminiferal assemblages are closely similar to two samples from harboøre and vorbasse, and the species composition found in a sample from an additional nearby interglacial deposit at hoven (fig. 1; semi-quantitative data) is also very close to those from kibæk. elphidium excavatum is dominant and ammonia beccarii is common in all the comparative records. age estimate the unusual presence of a lacustrine, as well as a marine sequence in one pleistocene series provides a minimum age of the bottom till and a maximum age of the marine deposit. the danish harreskovian interglacial pollen record can be correlated to the record at hunteburg in germany (hahne et al. 1994). a palaeomagnetic reversal at the base of the hunteburg series may indicate the presence of the brunhes–matuyama boundary, in which case the hunteburg interglacial series would correspond to marine isotope stage (mis) 19, almost 800 ka bp. the till at 39–45 m depth at kibæk would then belong to one of the till units identified in the lower part of the lind core (kronborg & odgaard 2004), deposited during one of the oldest pleistocene glaciations recorded from denmark. the marine unit represents a full interglacial period, but it was not possible to relate it to a specific interglacial. the deposit is, however, older that the eemian interglacial (mis 5e), because of the lack of lusitanian elements that charfig. 4. percentage distribution of selected foraminifera in the marine unit of the kibæk core (dgu 95.2875). note the different scales. fig 5. principal component analysis (pca) biplot of foraminiferal assemblages from kås hoved (knudsen et al. 2014), holstebro nord (kronborg et al. 2002), harboøre (knudsen 1987), vorbasse (k.l. knudsen, unpublished data), and kibæk (this study). a similar assemblage was also found in a sample from a core near hoven (dgu 103.1011; semiquantitative data in geus’ jupiter database). for locations see fig. 1. the percentage values were log-ratio transformed prior to the analysis. 17 18 19 20 21 0 0 0 0 0 0 0 0 elp hid ium ex cav atu m am moni a b ecc ari i bu lim ina m arg ina ta elp hid ium m age llan icu m elp hid ium alb ium bili cat um elp hid ium ge rth i hayn esi na ger mani ca fo ra m in ife ra p er 1 00 g 20 40 60 80 1010 10 101010 200 400 600 percentages number d ep th ( m ) inat a. beccarii h. orbiculare h. germanica e. gerthi e. albiumbilicatum e. magellanicum –1.0 kibæk samples species kås hoved 2 holstebro n vorbasse harboøre – 1 .0 1 .0 1.0 b. marginata e. margaritaceum e. excavatum 5050 acterise eemian assemblages in denmark (e.g. knudsen 1994). the assemblages from the marine unit are close to those related to holsteinian interglacial (mis 11) deposits in denmark (fig. 4), but the actual ages of these deposits are also uncertain. a stratigraphical correlation of the interglacial sediments at kås hoved with mis 11 was discussed by knudsen et al. (2014), but further development of optically stimulated luminescence (osl) dating or other absolute dating methods are needed for more exact age estimates of marine pre-eemian interglacial records in denmark. during the late early and the middle pleistocene, foraminiferal assemblages would be expected to be almost identical during similar marine environmental conditions. some of the danish records that are traditionally related to the holsteinian, could thus be from any warm interglacial period with high sea-level during the late early or middle pleistocene. the stratigraphical position of the marine unit in the kibæk core indicates a middle pleistocene age. it has been suggested that mis 11 was an exceptionally long and relatively warm interglacial (e.g. candy et al. 2014) and peak global sea level may have reached 8.0–11.5 m higher than today (chen et al. 2014). however, most temperature proxies for mis 11 give values similar to the holocene (candy et al. 2014), and according to bowen (2010) sea levels were close to the present sea level. because of the orbital similarity between mis 11 and mis 1, these are even used as analogues of warm intervals in climatic studies (berger & loutre 2003). the temperature indication of interglacial foraminiferal assemblages in denmark, which is related to the holsteinian, is also comparable with the holocene and present day temperatures in the region. it is remarkable that the middle pleistocene marine interglacial records at kibæk, vorbasse and hoven in central jylland are found several metres above sea level, whereas the harboøre and holstebro records in north-western jylland are found at more than 30 m b.s.l. a comparison of the core sites with the pre-quaternary surface map of denmark (binzer & stockmarr 1994) shows that kibæk, vorbasse and hoven are located in areas with a high-lying pre-quaternary surface, whereas harboøre and holstebro represent an area with a low-lying pre-quaternary surface. although glacial tectonics may have displaced some of the interglacial marine records the general elevational pattern of these sediments indicates that the central part of jylland may have experienced a tectonic uplift during the quaternary. references andersen, s.t. 1965: interglacialer og interstadialer i danmarks kvartær. meddelelser fra dansk geologisk forening 15, 486–506. berger, a. & loutre, m.-f. 2003: climate 400,000 years ago, a key to the future? american geophysical union geophysical monograph 137, 17–26. binzer, k. & stockmarr, j. 1994: geologisk kort over danmark. 1:500 000. prækvartæroverfladens højdeforhold. danmarks geologiske undersøgelse kortserie 44. bowen, d.q. 2010: sea level ~400 000 years ago (mis 11): analogue for present and future sea level? climate of the past 6, 19–29. candy, i., schreve, d.c., sherriff, j & tye, g.j. 2014: marine isotope stage 11: palaeoclimate, palaeoenvironments and its role as an analogue for the current interglacial. earth-science reviews 128, 18–51. chen, f., friedman, s., gertler, c.g., looney, j., o’connell, n., sierks, k. & mitrovica, j.x. 2014: refining estimates of polar ice volumes during the mis 11 interglacial using sea level records from south africa. journal of climate 27, 8740–8746. hahne, j., mengeling, h., merkt, j. & gramann, f. 1994: die hunteburg-warmzeit (“cromer-komplex”) und ablagerungen der elster-, saaleund weichsel-kaltzeit in der forschungsbohrung hunteburg ge 58 bei osnabrück. geologisches jahrbuch a134, 117–166. knudsen, k.l. 1987: elsterian-holsteinian foraminiferal stratigraphy in the north jutland and kattegat areas, denmark. boreas 16, 359– 368. knudsen, k.l. 1994: the marine quaternary in denmark: a review of new evidence from glacial-interglacial studies. bulletin of the geological society of denmark 41, 203–218. knudsen, k.l., ditlefsen, c., penney, d.n., kristensen, p., kronborg, c. & eiríksson, j. 2014: elsterian-holsteinian deposits at kås hoved, northern denmark: sediments, foraminifera, ostracods and stable isotopes. boreas 43, 251–271. kronborg, c. & odgaard, b.v. 2004. nyt om danmarks ældste kvartære aflejringer. dgf kvartærgeologisk møde november 2004. geologisk tidsskrift 2004(2), 23–24. kronborg, c., nielsen, o.b., sørensen, j. & kragelund, a. 2002: ringk øbing amt, holstebro nord, boring dgu nr. 64.1248. geologisk institut, aarhus universitet, report 02rk–01, 39 pp. authors’ addresses b.v.o. & k.l.k., department of geoscience, aarhus university, høegh-guldbergs gade 2, dk-8000 aarhus c, denmark. e-mail: bvo@geo.au.dk o.b. & h.j.g., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. geological survey of denmark and greenland bulletin 20, 2010, 9–14 9 the former director of grønlands geologiske undersøgelse (ggu; the geological survey of greenland), knud ellitsgaard-rasmussen, died on 1 december 2009, 86 years old. ellitsgaard was for many years a driving force in the build-up of ggu and became, as a relatively young geologist, its leader in 1956. in 1966 he was formally appointed director and remained in this position until his retirement in 1983. with ellitsgaard as director for 27 years, ggu developed from a small institute with a small permanent staff, into an internationally oriented research institute with a staff of more than c. 120, almost half of which were geoscientists. the survey’s activities were based on an integrated cooperation between ggu personnel and a very large group of external geoscientists who came both from danish universities and from international earth science institutes, mainly from great britain, holland, switzerland and scandinavia, to take part in ggu’s expeditions to greenland. every field season during the 1970s and 1980s, ggu sent between 100 and 150 participants to greenland. these carried out geological investigations throughout the immense country with emphasis on basic research and geological mapping, but with a gradually increasing focus on economic geology. the financial support for these many activities was a combination of government grants and grants for specific projects provided by the danish natural science research council and the ministry for trade (after 1979 the ministry for energy) as well as by private funds. during ellitsgaard’s directorship ggu’s budget grew several fold from a few million danish kroner a year to almost 50 million kroner in 1982. on 1 january 1984 ellitsgaard was succeeded as director by martin ghisler. on 1 june 1995 ggu was amalgamated with danmarks geologiske undersøgelse (dgu; the geological survey of denmark) to become the present geological survey of denmark and greenland (geus). thus ellitsgaard’s achievements over the years were entirely related to greenland and had no relation to the very wide range of activities now undertaken by geus. knud ellitsgaard-rasmussen originally trained as a joiner before he opted for a higher education and started studying geology, completing his studies in 1952 with a master’s degree from the university of copenhagen. while still a student he became involved in greenland, and in his first field season in 1946 he partook as an assistant in the initial mapping of the precambrian basement between nuuk (then godthaab) and disko bugt. in 1948 one of ellitsgaard’s tasks was to undertake a detailed study of a small group of islands with archaean low-metamorphic supracrustal rocks c. 10 km north-east of aasiaat (formerly egedesminde). the results were published in 1954, having already formed the basis of his master’s thesis for which he was awarded a gold medal from the university of copenhagen. in 1949–50 ellitsgaard partook in the danish peary land expedition under the leadership of the archaeologist eigil knuth. this meant staying in greenland for about a year and wintering in high-arctic conditions in the very desolate and isolated southern part of peary land in north greenland. taking advantage of winter conditions ellitsgaard, accompanied by a greenland sledge driver, travelled by dog sledge through the virtually unknown, northernmost part obituary: knud ellitsgaard-rasmussen 23 june 1923 – 1 december 2009 niels henriksen and t. christopher r. pulvertaft k. ellitsgaard-rasmussen, c. 2004. © geus, 2010. geological survey of denmark and greenland bulletin 20, 9–14. open access: www.geus.dk/publications/bull 1010 of peary land, mapping the deformed sedimentary rocks in the palaeozoic ellesmerian fold belt. the results of this reconnaissance survey were published in 1955. after receiving his master’s degree ellitsgaard was employed as a scientific assistant at the mineralogical-geological institute of the university of copenhagen. although at once involved in teaching, he continued to work in greenland together with colleagues from the university. establishing ggu from 1946 to 1956 after the second world war the government decided to initiate systematic geological investigations in greenland by setting up ggu, and an advisory committee was established for the coming geological survey. the committee came to consist of three geologists (professors arne noe-nygaard and alfred rosenkrantz and the director of the geological survey of denmark, hilmar ødum) together with the head of the greenland administration under the prime minister’s office. the focus in the first phase should be on geological mapping and the provision of the necessary geological expertise by means of training and cooperation with danish geologists at the university and mineralogical museum. the latter was achieved thanks to the efforts of noe-nygaard, who also, together with hans ramberg, initiated mapping of the precambrian basement in southern west greenland. studies of the cretaceous–tertiary sediments and basalts in central west greenland were led by rosenkrantz. several younger geologists and students were incorporated in the work, and ellitsgaard took part in organising the work as well as participating in the field work. a section was established for providing expeditions with the necessary equipment for field work (tents, sleeping bags, provisions etc.). two motor cutters transported field parties along the coast and into the fjords of greenland. during this period 20–30 persons participated each summer in the field work in west greenland, divided between precambrian basement areas and the cretaceous– tertiary of the disko–nuussuaq area. later more cutters were acquired as the number of parties in the field grew. as the scope of activities in greenland increased, the advisory committee wanted to be relieved of the day-to-day running of ggu and decided that ellitsgaard should be appointed its leader, first for a trial period of three years. this arrangement was extended stepwise until 1966 when he was formally appointed director. he continued for a period to ply his research interests by working in the spectrometer laboratory, mainly analysing samples collected in the west greenland basement areas, and he was lecturer in economic geology from 1962 to 1967. ggu was from the start an integrated part of the geological milieu at the university and the mineralogical museum (later renamed the geological museum), and ggu was allocated rooms together with university and museum geologists in the complex at øster voldgade 5–7 in copenhagen. at the same time the various institutes shared laboratories which were mainly used for analysis of greenland material. ggu’s tasks from 1956 to 1964 in 1956 the only systematic geological mapping that had been carried out in west greenland was the coastal reconnaissance of the area between 63°45́ n and 69°n by noenyegaard and ramberg. a very generalised map of southern west greenland had been published in ussing (1912), but otherwise the vast area of precambrian basement from kap farvel to thule was unmapped. consequently the advisory committee decided that the first priority of future geological activities should be preparation of geological maps. this priority became the guiding line for formulating ggu’s work programmes and was adhered to by ellitsgaard until his retirement in 1983. the new mapping campaign began in 1956 in the area around the cryolite mine at ivituut in southern west greenland, the choice of area being motivated by the hope of finding new deposits of cryolite, a mineral that had provided a substantial return to society since 1858. mapping was carried out at scale 1:20 000 with a view to publishing maps at scale 1:100 000. systematic mapping was extended to southern west greenland ellitsgaard in peary land, 1950. 11 while at the same time detailed studies of the ilímaussaq intrusion were initiated in cooperation with the atomic energy commission and copenhagen university. this intrusion hosts uranium-enriched rocks and therefore was seen as a possible source of uranium that could be used as fuel if a nuclear power station was to be built in denmark. the systematic mapping was carried out by geologists working in two-man teams from tents and supplied at regular intervals from two bases – one at ivituut and the other at dyrnæs near narsaq. ellitsgaard took part in the work in greenland and spent the greater part of several summers organising the work from dyrnæs. at first transport was provided entirely by ggu’s own motor cutters, but boats could only service camps at the coast, making it difficult for geologists to map more remote inland areas. ellitsgaard soon realised that the mapping could be carried out far more rationally with helicopter support, which at that time had seen limited use in greenland. with great perseverance ellitsgaard succeeded in obtaining funds for the purchase of two bell 47-j helicopters that together had sufficient capacity to transport two geologists and their entire camp. the 1958 season was the first with these helicopters in operation, and the number of participants in ggu’s campaign in southern west greenland was c. 70. this strategy, with two-man teams spread over a wide area and supported by helicopters and cutters operating from ggu’s own base, was continued and refined in the coming years. as mapping progressed ggu moved its base progressively northwards, and by the beginning of the 1980s this mapping had made it possible to compile 1:100 000 geological maps of the greater part of west greenland south of nuuk. in the 1960s systematic mapping was also carried out in west greenland north of nuussuaq where loose blocks of rich lead-zinc ore had been found in the 1930s. these blocks turned out to have fallen from a very rich lead-zinc deposit which was mined from 1973 to 1990 (black angel mine). due to the severe alpine terrain, much of this area was inaccessible on foot, necessitating a change in tactics. however, ellitsgaard realised that the solution in areas of this kind was to make intensive use of aerial photographs. consequently ggu took on an experienced geologist who was accustomed to mapping combined with photogeological interpretation, and 1:100 000 maps were also compiled in this area. later photogeological interpretation became an essential element in the mapping of north and north-east greenland. the preparation of 1:100 000 scale map sheets was accompanied by detailed studies of the material collected during the mapping, and by interpretation of the results. one of ellitsgaard’s great merits was to ensure from the start that there should be a scientific bonus from the mapping. to this end a large number of external, largely foreign, specialists were drawn into the work, cooperating closely with ggu’s own geologists. ggu’s extended activities by the middle of the 1960s it became evident that, with the speed at which mapping was progressing with the strategy then in use, it would be several decades before all greenland was covered by 1:100 000 geological maps. a new goal for ggu was therefore formulated. by means of reconnaissance mapping, regional maps at scale 1:500 000 should be produced. ellitsgaard’s aim was that within five years the whole of west greenland would be covered by 1:500 000 maps, given that areas already covered by 1:100 000 maps would not require remapping. in the end, however, it was many more years before all west greenland was covered by four 1:500 000 geological maps. in later years the concept was extended to ggu’s south greenland team at dyrnæs, 1959. ellitsgaard is seated second from the left in the front row. 1212 other parts of greenland, and today the whole of greenland is covered by 1:500 000 geological map sheets. an important expansion of ggu’s activities took place in 1967 when ggu took over the geological mapping of east and north greenland. previously geological investigations in these regions had been the preserve of the ‘danish expeditions to north-east greenland’ under the leadership of lauge koch. field work in these very remote and often extremely rugged areas required a much more intensive use of aeroplanes and helicopters than work in west greenland. this new activity substantially increased the scope of ggu’s tasks, as mapping in west greenland was to continue with the same intensity as before. ellitsgaard mobilised his persuasive talents and in 1968 secured funds enabling ggu to charter a polar vessel, not only to transport the expedition to and from east greenland, but also to act as a floating base for operations throughout the season. two helicopter landing platforms were mounted on the ship. this strategy was used for three seasons, after which operations were based entirely on tent base camps on land. the expedition with all its goods was transported to and from greenland in a large aircraft, and small stol aircraft and helicopters provided the daily transport of the mapping operation. as mapping progressed, base camps were moved according to logistic needs. by the early 1970s ggu’s combined operations in west and east greenland included 120–130 persons, five to six chartered helicopters and stol aircraft, and five cutters. there was plenty for ggu’s director to see to! ellitsgaard was always aware of the advantages to ggu of allowing geologists to partake in activities led by other organisations. thus in north greenland ggu personnel participated in the geological survey of canada’s operation grant land in 1965–66 and in a reconnaissance in peary land in 1969 under the auspices of a british joint services expedition. thanks to these reconnaissance activities and its own pilot studies in the 1970s, ggu was well prepared when systematic mapping of north greenland was started in 1978. by 1970 ggu had made sufficient progress to compile the first ever geological map of the whole of greenland at scale 1:2 500 000. this was not superseded until 1995 when a completely revised map at this scale was published. unlike the 1970 map, this later map included offshore geology. publication of the first geological map was followed up in 1971 by the publication of the first map of the quaternary geology of greenland, also at scale 1:2 500 000. although ggu’s primary task was geological mapping, this was not its only activity. from the middle of the 1950s ggu was involved in economic geology, carrying out pilot studies for mining companies. in the 1960s the scope of ggu’s glaciological studies was expanded to assessing possible sites for establishing hydroelectric power stations, with the result that the first hydropower station in greenland could be inaugurated in 1993. in the late 1960s the oil industry began to show an interest in west greenland, particularly offshore southern west greenland. ellitsgaard soon realised that ggu would need to react to this development, and a department for oil geology was established with three functions: (1) acquiring oil-geologically relevant data from onshore sedimentary basins that could provide analogues one of ggu’s bell helicopters, 1958. ellitsgaard at dyrnæs, c. 1960. 13 to what can be expected to occur in offshore basins, (2) acquiring geophysical data offshore in relatively ice-free areas that were regarded as having a potential for hydrocarbon deposits and (3) advising the then ministry for greenland in assessing applications for petroleum exploration licences. the first licensing round covering blocks offshore southern west greenland was opened in 1974. five exploration wells were drilled in 1976–77, but all were declared dry and exploration in this region was not resumed until 1991. ggu’s offshore activities made a modest start in 1972 with a shallow seismic reconnaissance of the west greenland shelf and fjords between 68° and 73°n. following the energy crisis in 1973 the natural science research council and the ministry for trade (after 1979 the ministry for energy) made funds available for energy-related projects. ggu secured its share of these funds, and in 1978 ggu ‘went offshore’ in earnest. with support from the danish energy authority under the ministry for trade, extended shallow marine surveys of the west greenland shelf were carried out between 62° and 68°n. in 1979 ggu turned its attention to the east greenland shelf which, prior to the opening of the north atlantic, lay close to the major oil-producing areas of the north sea and west norwegian shelf. to start with, an aeromagnetic survey of the region was carried out in 1979 (project eastmar). with additional support from the eec this was followed up by a reflection seismic survey (north atlantic d) in 1980–82. other energy-related projects were a detailed mapping and evaluation of the coal deposits on nuussuaq and an extensive uranium-prospecting project in south greenland (project syduran), the latter inspired by reports of uranium occurrences in eastern labrador in a geological setting similar to that in south greenland. from embryonic survey to formal state survey as soon as ellitsgaard had been appointed leader of the institute now officially called grønlands geologiske undersøgelse, the advisory committee wanted to hand over the full responsibility to ellitsgaard, and ggu to be given a formal legal status. this suggestion was well received. however, the ministry for greenland also wanted a mining law for greenland to be drafted and, as ggu would have a role to play in connection with mining activities, it was decided to hand these tasks over to a mining law commission. two laws – the mining law and the law for ggu – were finally passed in the danish parliament (folketinget) in 1965 and ggu became a directorate under the ministry for greenland. shortly afterwards ellitsgaard was appointed ggu’s first director after functioning as such for ten years with uncertain tenure. after this the committee for ggu dissolved itself (noe-nygaard 1986). external cooperation when work began in greenland in 1946 there were no geologists in denmark with expertise in basement geology. this situation was remedied, partly thanks to a determined effort by professor arne noe-nygaard at the university of copenhagen, and partly by recruiting foreign geologists with experience in this field. with this policy the seed was sown for ggu later to become an internationally oriented institute, where geological mapping and research were carried out by danish and foreign geologists employed at ggu in cooperation with a large number of research students and geologists based in foreign universities. the latter were attracted by the fantastic natural conditions for geological research offered by greenland and by the very favourable working conditions offered by ggu. thus an extensive ggu’s base camp at ‘mellembygd’, near paamiut (formerly frederikshåb), 1964. ggu’s floating base in scoresby sund in 1968. the ship martin karlsen is the vessel formerly named kista dan. 1414 international network of partners was built up, which raised the scientific standard in ggu to a level comparable to that of many leading international research institutes. ggu and greenland geology also became more widely known in the western world when a review of all aspects of greenland geology was published in 1976 (escher & watt 1976). in a review of this book john sutton, then professor of geology at imperial college, london, wrote: “the survey collaborates with more than fifty universities and institutes, and moreover does so with a generosity and openness which has attracted able scientists from many countries. the heart of the undertaking lies, however, in denmark. it is danish resources and leadership that have brought the geological survey of greenland to its present eminence.” (sutton 1976, p. 815). a measure of ggu’s high standard during this period is the fact that more than 25 of the then young geologists later became full professors at danish and foreign universities. honours throughout his career ellitsgaard partook in the work of several committees, commissions and scientific societies. notably ellitsgaard was a member of the commission for scientific research in greenland from 1965 to 1983, a member of the danish academy for technical sciences from 1967 to 1997, and in 1974 he was elected a member of the royal danish academy of sciences and letters. he was also a member of the board of governors for nordisk mineselskab a/s and arktisk minekompagni a/s. from 1981 to 1986 he was a member of greenland home rule’s national park council. in 1976 ellitsgaard received the egede medal of the royal danish geographical society for his contribution to geological and geographical research in greenland. the high esteem in which ellitsgaard was held in foreign geological institutes was shown when he was awarded an honorary doctorate of science at the university of exeter (uk) in 1984 and furthermore elected an honorary fellow of the geological society of london. final words during the 27 years when he bore the main responsibility for the administration and development of ggu, ellitsgaard succeeded in building up an organisation that became widely known in international geological circles and had a great influence on the coming development of economic geology in greenland. ellitsgaard had his principles and views and at times ran into political and administrative problems. however, he tackled these difficulties without ggu’s personnel being affected by them. he has written his personal perception of the embryonic days of ggu in an internal survey report (ellitsgaard-rasmussen 1996). ellitsgaard had many good years as ggu’s director, and his staff remember him as a friendly and approachable leader, although heavily involved in administrative duties. his philosophy for the daily leadership of the institute was that his staff would perform best if given a large degree of individual freedom and responsibility, a philosophy the success of which can be seen in ggu’s substantial production of scientific papers and map sheets. all honour to his name. references ellitsgaard-rasmussen, k. 1996: en stjerne fødes. beretning om ggu’s tilblivelse. danmarks og grønlands geologiske undersøgelse rapport 1996/102, 76 pp. escher, a. & watt, w.s. (eds) 1976: geology of greenland, 603 pp. copenhagen: geological survey of greenland. noe-nygaard, a. 1986: til knud ellitsgaard-rasmussen (in danish, with english summary). rapport grønlands geologiske undersøgelse 128, 5–11. sutton, j. 1976: geological survey of greenland. nature 264, 815 only. ussing, n.v. 1912: geology of the country around julianehaab, greenland. meddelelser om grønland 38, 376 pp. k. ellitsgaard-rasmussen receiving the honorary degree of doctor of science at the university of exeter, u.k., in 1984. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: nielshen@mail.dk geological survey of denmark and greenland bulletin 20, 2010, 95–98 95 beijing jinan tianjin tangshan tianjin greengen power plant dagang oilfield complex shengli oilfield complex kailuan coalfields jiyang aquifers (huimin sub-basin) 50 km 116°e 118°e 40°n 38°n china pipelines, fluid conveyed gas gas under construction or planned oil oil under construction or planned other pipelines under construction or planned aquifers oilfields kailuan coalfields status projected proved city fig. 1. map of the study area in eastern china showing co2 sources, proposed pipeline network and potential storage sites. based on data from the energy, environment and economy research institute, tsinghua university; institute of geology and geophysics, chinese academy of sciences; china university of mining and technology; research institute of petroleum exploration and development, petrochina and the china university of petroleum (cup). the outline of the shengli oilfield complex and the pipeline data are from ‘energy map of china 2008’, © the petroleum economist ltd, london. © british geological survey. british geological survey produced the gis map. the challenge of climate change demands reduction in global co2 mission. carbon dioxide capture and storage (ccs) technology can be used to trap and store carbon dioxide gas emitted by coal-burning plants and this can reduce the world’s total co2 emission by about one quarter by 2050 (iea 2008, 2009; ipcc 2005). experience from the storage sites of sleipner in the norwegian north sea, salah in algeria, nagaoka in japan, frio in usa and other sites shows that geological structures can safely accommodate co2 produced and captured from large co2 point sources. ccs is regarded as a technology that will make power generation from coal sustainable, based on cost-effective co2 capture, transport and safe geological storage of the released co2. china has large coal reserves (delaquil et al. 2003), and is not about to give up on this reliable source of fossil fuel. hence a large production of co2 can be expected to continue for many years. china also has a large theoretical geological carbon dioxide storage capacity in onshore areas with deep saline formations (dahowski et al. 2009). in an extensive collaborative research effort between chinese and european scientists, the coach project (cooperation action within ccs china-eu) was successful in building the expertise, developing the capture technologies and mapping transportation routes for co2, and it produced two scenarios for geological storage of co2 in china. the aim of the coach project was to initiate a durable cooperation between europe and china in response to china’s rapidly growing energy demand. the project ran from november 2006 to october 2009 and was set up and funded by the european commission under the memorandum of understanding on near zero emissions coal, to build demonstration plants in china. twenty partners consisting of eight chinese and twelve european partners evaluated the feasibility of establishing ccs in china (coach 2009). coach had four work packages dealing with (1) knowledge sharing and capacity building, (2) capture technologies, (3) permanent geological storage of co2 and (4) recommendations and guidelines for implementation. three tasks were carried out under the potential for permanent geological storage of co2 in china: the coach project niels e. poulsen © geus, 2010. geological survey of denmark and greenland bulletin 20, 95–98. open access: www.geus.dk/publications/bull 9696 third work package: (a) capacity estimates at regional level, (b) mapping of the geology and emission point sources and (c) improving methods for storage capacity assessment and site selection criteria. the geological survey of denmark and greenland and tsinghua university in beijing shared the leadership of the third work package. this short article presents the results of the work conducted on the potential for geological storage of co2 in china. background and methods aims of the carbon sequestration leadership forum the aim of co2 storage is the permanent removal of co2 from the atmosphere. the european union has supported current research on co2 capture and storage methods for more than a decade, with emphasis on capture techniques, transport and geological storage. the results of the research on geological storage are summarised in a comprehensive manual by chadwick et al. (2008). internationally recognised standards for capacity assessments were established by the carbon sequestration leadership forum (cslf) in 2004–2005 and a task force on capacity estimation standards has been active since presenting comprehensive definitions, concepts and methods (bachu et al. (2007a, b). these capacity standards were reviewed for the coach project by poulsen et al. (2009) and were used for the work on permanent co2 storage estimates in china (zeng et al. 2009). comparison of methods various methods are available for calculation of co2 storage capacity in a geological environment (koide et al. 1992, 1995; tanaka et al. 1995; shafeen et al. 2004) . the methods used in the coach project (poulsen et al. 2009) were based on bachu et al. (2007a, b) and used in the coach database to estimate the storage capacity of hydrocarbon fields. estimates made by the china university of petroleum applied tanaka et al.’s (1995) method for computing the storage capacity in the shengli oilfield complex (zeng et al. 2009). the two methods proposed by the cslf task force and tanaka et al. (1995) are basically identical in their approach. both methods are based on a volumetric approach and are applicable to site, regional and basin-scale co2 storage capacity estimates. both can be considered as ‘simple’ equation models, which try to calculate an ‘approximation’ of a possible storage capacity. the methods gave almost identical results when applied to the shengli oilfield complex (table 1). there are, however, some differences in the approach to co2 behaviour in the storage site. the cslf method works with replacement of oil, gas or formation water but does not incorporate dissolution of co2 in formation water. the method of tanaka et al. (1995), on the other hand, operates with a free phase of co2 and takes into account dissolution of co2 in the formation water, but it does not considerer the time period needed for the dissolution (poulsen et al. 2009). long term behaviour of co2 in a storage site the long term behaviour of co2 in a storage site depends on (1) a number of reservoir parameters (temperature, pressure, capillary pressure, porosity, permeability, and the cap rock permeability and capillary entry pressure), (2) the co2 composition, (3) the formation water and (4) time (chadwick et al. 2008). the solubility of co2 in formation water varies with salinity, temperature and pressure of the formation water (the brine). the dissolution of co2 in pure water increases with increasing pressure (and thus increasing depth) up to approximately 7 mpa. on the other hand, the co2 solubility in a brine decreases with increasing temperature and salinity and thus in most cases decreases with depth (bachu & adams 2003). the fig. 2. an example of a shengli oilfield production site. 97 result is that in general, the solubility of co2 in the brine decreases with increasing salinity (shafeen et al. 2004). the buoyancy of injected supercritical co2 leads to an upward gravity-driven flow of co2 towards the top of the formation where it forms a plume below the cap rock. co2 (liquid or supercritical) and water are immiscible, but co2 can dissolve to a certain extent in water. due to the slow solubility of co2 in brine, a large volume of brine is necessary to dissolve a given amount of co2. the density of the brine increases with increasing co2 dissolution and a downward gravity-driven flow will be induced by the increased density of the co2-saturated brine. on the initiation of storage, before the plume of saturated brine has reached the bottom, the overall dissolution rate is essentially constant due to rapid convective overturn (ennis-king & paterson 2007). at a later stage during storage the saturated brine forms a gravity current propagating outwards from the co2 source. activities and results the main purpose of the coach project was to prepare the way for co2 capture and storage in china. in order to achieve this, the coach partners developed an integrated gasification combined cycle capture technique. this is a coal-based energy system with hydrogen production using coal gasification, electricity generation from a combined cycle hydrogen turbine and fuel cell system, and capture of the co2. the partners have mapped emission sources and investigated potential co2 storage sites in eastern china (fig. 1, table 1). the storage potential of the selected sites was evaluated using published data or data provided by the research institute of petroleum (petrochina). particular oilfields, saline aquifers and un-exploitable coal beds were investigated. several test sites are available in some of the oilfields. the storage potential in oilfields is 10–500 mt, (pilot scale level; fig. 1, table 1). following this, a co2 transport infrastructure based on connecting co2 sources and storage sites by pipeline or ship has been suggested (fig. 1; table 1). the saline jiyang aquifers in the huimin sub-basin show storage capture at an industrial scale (around 50 gt; fig. 1, table 1), but further geological investigations are required. the security of energy supply is a key consideration in china, and enhanced oil recovery (eor) could be an option. some of the oilfields in the dagang and shengli oilfield complexes may be suitable for an enhanced oil recovery pilot project. injecting co2 into oilfields approaching depletion will not only store co2, but may also enhance or prolong oil recovery (coach 2009). the coals of the kailuan coalfield have low permeability and probably low injectivity, but a high theoretical ability to adsorb co2 (fig. 1, table 1). in general, however, the storage capacity in coal seams is uncertain. on the other hand, it has been demonstrated that injection of co2 into coal beds can lead to methane production (enhanced coal bed methane recovery; yu et al. 2007). at the same time it is a very attractive option for geological co2 storage as co2 is strongly absorbed onto the coal. two scenarios for possible co2 capture and storage demonstration projects have been proposed by work package 4, based on the mapping of emission point sources, geology, and capacity estimates by work package 3 together with economic analyses. the first scenario is for a pilot scale site with 0.1–1 mt co2/year stored in the dagang or shengli oilfield complexes. the second scenario is intended for industrialscale storage at 2–3 mt co2/year, which could be accommodated in the shengli oilfield complex or potentially in the saline formations in the huimin sub-basin. the pilot scale scenarios focus initially on enhanced oil recovery for storage where this is feasible. the large-scale option could begin with enhanced oil recovery but would need to switch to saline 472 mt using cslf methodology and 463 mt using cup method table 1. summary of geological sites assessed for geological storage of co2 after zeng et al. (2009) storage site capacity injectivity seal dagang oilfield complex selected 7 fields 22 mt largest gangdong field 10 mt 1000 md some compartmentalisation by faulting and stratigraphy mudstones shengli oilfield complex 1000–2500 md some compartmentalisation by faulting and stratigraphy lower jurassic mudstones huimin sag aquifers (jiyang) for huimin sub-basin 50 gt for selected troughs in sub-basin 0.7 gt permeability around 1600 md in neighbouring oilfields mudstones of minghuanzhen fm kailuan coalfield 504 gt adsorbed onto coal and 38 100 mt void capacity permeability generally low 3.7 md in taiyuan formation and 0.1 md in shanxi and xiashihezi fm mudstones 9898 aquifer storage once the potential reservoir and sealing formations have been adequately investigated. both scenarios are based on capture of co2 from the tianjin greengen power plant (coach 2009). final remarks in 2005 construction began of the coal-based tianjin greengen power plant (fig. 1) and electricity production started in 2009. it will be the first near-zero emission power plant in china. research over the next decade is expected to develop and demonstrate the efficiency of coal-based power generation, mostly by recycling energy lost in the process. the goal is to achieve sustainability of coal-based power generation. the project concludes that there is significant potential to develop carbon dioxide capture and storage technologies in china and to make major reductions in co2 emissions over the next century. experience from the storage sites sleipner in the norwegian north sea, in salah in algeria, nagaoka in japan, frio in usa and other sites shows that geological structures can safely accommodate co2 produced and captured from large point sources. thus, geological storage of co2 can contribute considerably to the reduction of co2 emission in china and other countries. acknowledgements coach was funded as part of the 6th framework programme for research by the european commission (project no. 038966). nikki smith from the british geological survey is thanked for producing the map used in fig. 1. references bachu, s., & adams, j. j. 2003: sequestration of co2 in geological media in response to climate change: capacity of deep saline aquifers to sequester co2 in solution. energy conversion and management 44, 3151–3175. bachu, s., bonijoly, d., bradshaw, j., burruss, r., christensen, n.p. holloway, s., & mathiassen, o.m. 2007a: estimation of co2 storage capacity in geological media – phase 2. work under the auspices of the carbon sequestration leadership forum (www.cslforum.org). final report from the task force for review and identification of standards for co2 storage capacity estimation, 43 pp. washington: carbon sequestration leadership forum. bachu, s., bonijoly, d., bradshaw, j., burruss, r., holloway, s., christensen, n.p. & mathiassen, o.m. 2007b: co2 storage capacity estimation: methodology and gaps. international journal of greenhouse gas control 1, 430–443. chadwick, a., arts, r., bernstone, c., may, f., thibeau, s. & zweigel, p. (eds) 2008: best practice for the storage of co2 in saline aquifers – observations and guidelines from the sacs and co2store projects. british geological survey occasional publication 14, 267 pp. coach 2009: project n° 038966: coach, cooperation action within ccs china-eu, executive report, 38 pp. dahowski, r.t., li, x., davidson, c.l., wei, n., dooley, j.j. & gentile, r.h. 2009: a preliminary cost curve assessment of carbon dioxide capture and storage potential in china. energy procedia 1, 2849–2856. delaquil, o., wenying, c., & larson, e.d. 2003: modeling china’s energy future. energy for sustainable development 7, 40–56. ennis-king, j. & paterson, l. 2007: coupling of geochemical reactions and convective mixing in the long-term geological storage of carbon dioxide. international journal of greenhouse gas control 1, 86–93. iea (international energy agency) 2008: energy technology perspectives: scenarios and strategies to 2050, 650 pp. paris, france. iea (international energy agency) 2009: technology roadmap. wind energy, 52 pp. paris: international energy agency. koide, h., tazaki, y., noguchi, y., nakayama, s., iijima, m., ito, k., & shindo, y. 1992: subterranean containment and long term storage of carbon dioxide in unused aquifers and in depleted natural gas reservoirs. energy conversion management 33, 619–626. koide, h., takahashi, m., tsukamoto. h. & shindo, y. 1995: self-trapping mechanism of carbon dioxide in aquifer disposal. energy conversion management 36, 505–508. metz, b. et al. (eds) 2005: carbon dioxide capture and storage. ipcc 2005, 431 pp. cambridge university press. poulsen, n.e., chen, w., dai, s., ding, g., kirk, k., li, m., zeng, r., vangkilde-pedersen, t., vincent, c.j. & vosgerau, h.j. 2009: d3.3. improving methodologies for storage capacity assessment and site selection criteria. eu project no. 038966. coach work package 3 report. eu deliverable d3.3, 45 pp. eu coach project, brussels. shafeen, a., croiset, e., douglas, p.l. & chatzis, i. 2004: co2 sequestration in ontario, canada. part i: storage evaluation of potential reservoirs. energy conversion and management 45, 2645–2659. tanaka, s., koide, h. & sasagawa, a. 1994: possibility of underground co2 sequestration in japan. energy conversion and management 36, 527–530. yu, h., zhou, g., fan, w. & ye, j. 2007: predicted co2 enhanced coalbed methane recovery and co2 sequestration in china. international journal of coal geology 71, 345–357. zeng, r., li m., dai, s., zhang, b., ding, g. & vincent, c. 2009: assessment of co2 storage potential in the dagang oilfield, shengli oilfield and kailuan coalfield. coach work package 3 report. eu deliverable d3.1, 45 pp. eu coach project, brussels. author’s address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: nep@geus.dk geological survey of denmark and greenland bulletin 20, 2010, 15–18 15 the danish oil and gas production mainly comes from fields with chalk reservoirs of late cretaceous (maastrichtian) and early paleocene (danian) ages located in the southern part of the danish central graben in the north sea. the area is mature with respect to exploration with most chalk fields located in structural traps known since the 1970s. however, the discovery by mærsk oil and gas a/s of the large nonstructurally and dynamically trapped oil accumulation of the halfdan field in 1999 north-west of the dan field (e.g. albrechtsen et al. 2001) triggered renewed exploration interest. this led to acquisition of new high quality 3-d seismic data that considerably enhanced imaging of different depositional features within the chalk group. parallel to the endeavours by the operator to locate additional non-structural traps in porous chalk, the geological survey of denmark and greenland took advantage of the new data to unravel basin development by combining 3-d seismic interpretation of a large number of seismic markers, well log correlations and 2-d seismic inversion for prediction of the distribution of porous intervals in the chalk group. part of this study is presented by abramovitz et al. (in press). in the present paper we focus on aspects of the general structural development during the late cretaceous as illustrated by semi-regional time-isochore maps. the chalk group has been divided into two seismically mappable units (a cenomanian–campanian lower chalk unit and a maastrichtian–danian upper chalk unit) separated by a distinct basin-wide unconformity. study area and database the mapped area comprises the southern salt dome province located in the southern part of the danish central graben and adjacent parts of the ringkøbing–fyn high (fig. 1). the seismic database used involves three separate 3-d data sets. the bulk of the area is covered by the 1150 km2 kraka extension survey, which was acquired by mærsk in 2000. to the south and east it is supplemented by data belonging to the fugro multi-client entenschnabel 2002 survey and to the north-east by part of the mærsk contiguous area 3-d from 1995. in addition, data from smaller surveys of different vintages were used in order to cover parts of the coffee soil fault that separates the graben area from the platform. geological setting deposition of the upper cretaceous – danian chalk group in the study area took place during a phase of regional postrift subsidence following late jurassic rifting. this period was marked by high sea level, high seawater temperatures and a peak in production of organic matter. the late cretaceous regional subsidence was modified by movements of zechstein salt and punctuated by widespread inversion in the form of compression along old extensional fault trends, resulting in flexuring and folding of basin infill (vejbæk & andersen 2002). this resulted in the development of areas with bathymetric elevations and formation of local depocentres in the intervening lows. the structural movements gave rise to a number of unconformities easily recognised as truncation and onlap surfaces on seismic profiles and to stratigraphic hiatuses in wells. the chalk in the area is pelagic in origin and formed from settling of calcareous nannoplankton remains (coccoliths) and it consists generally of 95–99% calcite. after deposition the pelagic chalk was subjected to redistribution late cretaceous basin development of the southern danish central graben finn jakobsen and claus andersen fig. 1. map of the danish central graben region, showing the major structural elements and the location of the study area. © geus, 2010. geological survey of denmark and greenland bulletin 20, 15–18. open access: www.geus.dk/publications/bull norwegian–danish basin c offee soil fault ringkøbing–fyn h igh 4°e 5°e 56°n southern salt dome province central graben norway 25 km uk germany national border mapped area mapped areauk nl n dk 200 km g 1616 by various processes. these include downslope mass-flow movements arising from slope instability caused by syndepositional tectonic movements, or along-slope bottom currents that modified the seafloor forming channels, drifts, ridges and mounds. the interaction of downslope and along-slope processes has been demonstrated by esmerode et al. (2008) in the lower part of the chalk section in the area immediately to the north of the study area. the overall structural style is illustrated by a timestructure map of the top chalk group (fig. 2) and a timeisochore map of the total chalk group (fig. 3). the latter shows very large thickness variations from less than 100 msec (< c. 200 m) on top of salt diapirs up to 650 msec (c. 1300 m) in the rim-syncline east of the dagmar field. reduced thicknesses have been mapped over the dan and kraka structures, which are caused by growth of underlying salt pillows. on the nnw–sse-trending ‘gorm–lola ridge’ extending southwards from the gorm field, the thinning of the chalk is caused by a combination of halokinesis and structural inversion. the asymmetric, nw–se-trending ‘igor–emma ridge’ with less than 100 msec of chalk on its central part is located along the coffee soil fault. the total chalk thickness is much larger on the stable ringkøbing–fyn high than on the adjacent inverted part of the graben. the time-structure map of fig. 2 shows the results of the late cretaceous tectonic movements combined with the effects of cenozoic continued inversion/subsidence concentrated in the former graben area and halokinesis. structural inversion affecting the top chalk group surface is less pronounced and is restricted to the development of the centrally located, gentle, nw–se-orientated ‘tyra–igor ridge’ trapping the tyra, tyra se and halfdan ne gas accumulations. the large, dynamically trapped halfdan oilfield is located on the south-western flank of this ridge. basin development in modern 3-d seismic data, the chalk group is often characterised by discontinuous reflection patterns indicating spatial variations in lithology and depositional mode, and it is possible to identify and map a number of unconformities. the most prominent is here termed the top hod unconformity (fig. 4), which is generally expressed as a strong positive reflection. the unconformity divides the chalk group into a lower part comprising the hod, blodøks and hidra formations and an upper part comprising the ekofisk and tor formations using the standard lithostratigraphic nomenclature for the central north sea (surlyk et al. 2003). new palaeontological age determinations of the olga-1x well, which is located on the flank of the kraka structure, suggest that the unconformity here is of latest campanian to early maastrichtian age (abramovitz et al. in press). the seismic profiles in fig. 4 are selected to illustrate reflection patterns and highlight aspects of the basin development. the northern profile in fig. 4a crosses the gentle ‘tyra–igor ridge’ that was mainly formed by inversion postdating chalk deposition. several generations of channel-like features cut into the top hod unconformity. furthermore, the lower part of the chalk is extensively disturbed on the eastern flank. esmerode et al. (2008) interpreted similar features as formed by interactions between downslope mass movements (slumps and slides) and along-slope currents. the uppermost part of the chalk is draped and rather uniform in thickness suggesting reduction in earlier bathymetric relief. 2700 2500 2300 2100 1900 1700 10 km msec profile 1 tyra tyra se gorm dagmar rolf nana-1xp olga-1 regnar kraka skjold profile 2 profile 3 halfdan ne c o ff ee s o il fa u lt dan halfdan 5°e 55°25´n fig. 2. time-structure map of the top chalk group with location of producing chalk fields, profiles and wells mentioned in the text. contours on the shallow salt diapirs with thin chalk are not shown. 10 km 500 400 300 200 100 0 msec 600 700 55°25´n c o ff ee s o il fa u lt ‘g o rm –lo la r idge’ ‘igor–em m a ridge’ 5°e fig. 3. time-isochore map of the total chalk group with names of major inversion ridges indicated. chalk group isochores over the salt diapirs are not shown. 17 the top chalk group surface on the eastern flank is cut by low-angle listric faults, indicating early post-danian slumping of poorly consolidated chalk caused by renewed inversion along the ridge. figure 4b illustrates the asymmetric ‘igor–emma ridge’ inversion anticline trapping the halfdan ne gasfield found in danian chalk. the anticline is a result of both intra-chalk and post-chalk compression movements and developed with a steep north-eastern limb adjacent to the reversed coffee soil fault. the lower part of the chalk is deeply eroded at the crest and the top hod unconformity merges with older truncation surfaces. the tor formation above onlaps the western flank. the thinning and condensation of the tor formation is associated with deterioration of reservoir properties in the upper maastrichtian tor formation chalk arresting up-dip oil migration from the halfdan field, which was discovered from the nana-1xp vertical well. figure 4c shows a w–e profile that crosses the coffee soil fault. the profile is located in the southern part of the study area that is characterised by an asymmetric basin development with thick lower chalk below the top hod unconformity in the central graben area to the west, and an eastward shift of depocentre above. the lower part of the tor formation is thick on the ringkøbing–fyn high. it gradually onlaps the unconformity indicating significant variations in time span of missing sections across this surface. fig. 4. a: profile 1 crosses the gentle ‘tyra–igor ridge’ mainly formed by inversion post-dating chalk deposition. several generations of channel-like features cut into the top hod unconformity. b: profile 2 illustrates the asymmetric ‘igor–emma ridge’ inversion anticline. the anticline is a result of both intra-chalk and post-chalk compression movements and developed with a steep north-eastern limb adjacent to the reversed coffee soil fault. c: profile 3 is located in the southern part of the study area that is characterised by an asymmetric basin development with an eastward shift in depocentres. yellow lines show faults. for location of the profiles see fig. 2. w e e coffee soil fault 3 km top chalk intra hod base chalk 3 km ne 4 km w coffee soil fault 1500 2000 2000 2500 2500 2000 tw o -w ay t ra ve l ti m e (m se c) 2500 b a c w top hod unconformityintra tor nana-1xp coffeecoffee soilsoil faultfault coffeecoffee soilsoil faultfault coffeecoffee soilsoil faultfault ‘tyra–igor ridge’‘tyra–igor ridge’ ‘igor–emma ridge’‘igor–emma ridge’ coffee soil fault coffee soil fault coffeecoffee soilsoil faultfault ‘tyra–igor ridge’‘tyra–igor ridge’ ‘igor–emma ridge’‘igor–emma ridge’ coffee soil fault ‘tyra–igor ridge’ ‘igor–emma ridge’ 1818 the general basin development during the late cretaceous and danian can further be visualised by comparing the semi-regional time-isochore maps of the lower and upper parts of the chalk group separated by the top hod unconformity. the lower chalk group interval shows large thickness variations with the main depocentre (up to 350 msec) found in the rim-syncline north and east of the dagmar salt diapir (fig. 5a). the ‘gorm–lola ridge’ separates it from a nw–se-orientated depocentre that covers the southern flank of the kraka structure. the lower chalk group interval is also thin over the dan field area, indicating structural growth caused by underlying halokinetic movements. a conspicuous feature is the asymmetric ‘igor–emma ridge’ inversion anticline along the reversed coffee soil fault where the lower chalk is thin and even seismically absent in the crestal part. the time-isochore map of the upper chalk group interval (fig. 5b) shows more gradual thickness variations. the most notable difference between the two units is the northeastward shift and widening of the nw–se-orientated depocentre. at the dan field only an insignificant thinning of the interval is mapped, which shows that halokinetic movements waned here during the maastrichtian and danian. the expression of both the ‘gorm–lola ridge’ and the ‘igor–emma ridge’ is less distinct. the latter is confined to the northern part of the mapped area and orientated slightly obliquely to the nnw–sse-trending segments of the coffee soil fault zone. by using a simple two-fold subdivision of the chalk group, the present paper illustrates the dynamic basin development with spatial changes in subsidence patterns and structural growth caused by halokinesis and structural inversion. a more detailed analysis of the basin development with an additional subdivision of the chalk group is in progress in combination with mapping of sedimentary features. based on the close relationship between erosion, re-deposition (deposition of reservoir chalk) and structural development, the basin development is used as a guide for prediction of hitherto unrecognised reservoir intervals. references abramovitz, t., andersen, c., jakobsen, f.c., kristensen, l. & sheldon, e. (in press): 3-d seismic mapping and porosity variation of intra-chalk units in the southern danish north sea. in: vining, b.a. (ed): petroleum geology: from mature basins to new frontiers. proceedings of the 7th petroleum geology conference. london: geological society. albrechtsen, t., andersen, s. j., dons, t., engstrøm, f., jørgensen, o. & sørensen, f.w. 2001: halfdan: developing non-structurally trapped oil in north sea chalk. paper spe 71322. new orleans, louisiana: society of petroleum engineers annual technical conference and exhibition. esmerode, e.v., lykke-andersen, h. & surlyk, f. 2008: interaction between bottom currents and slope failure in the late cretaceous of the southern danish central graben, north sea. journal of the geological society (london) 165, 55–72. surlyk, f., dons, t., clausen, c.k. & higham, j. 2003: upper cretaceous. in: evans, d. et al. (eds and co-ordinators): the millennium atlas: petroleum geology of central and northern north sea, 213–233. london: geological society. vejbæk, o.v. & andersen, c. 2002: post mid-cretaceous inversion tectonics in the danish central graben – regionally synchronous tectonic events? bulletin of the geological society of denmark 49, 129–144. 10 km10 km 250 200 150 100 50 0 300 350 400 a bmsec 250 200 150 100 50 0 300 350 400 msec dagmar kraka c o ff ee s o il fa u lt dan ‘g o rm – lo la r id ge’ ‘igor–em m a ridge’ c o ff ee s o il fa u lt dan ‘g o rm –lo la r idge’ ‘igor–emma ridge’ 5°e 5°e 55°25´n 55°25´n fig. 5. time-isochore maps. a: lower chalk group interval. b: upper chalk group interval. the isochores show shift of depocentres during the late cretaceous in a north-eastward direction. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: fj@geus.dk geological survey of denmark and greenland bulletin 42, 2018, 133-147 133 burial and exhumation history of the jameson land basin, east greenland, estimated from thermo chronological data from the blokelv-1 core paul f. green and peter japsen apatite fission-track analysis (afta) data in two upper jurassic core samples from the 231 m deep blokelv-1 borehole, jameson land, east greenland, combined with vitrinite reflectance data and regional afta data, define three palaeo-thermal episodes. we interpret localised early eocene (55– 50 ma) palaeotemperatures as representing localised early eocene heating related to intrusive activity whereas we interpret late eocene (40–35 ma) and late miocene (c. 10 ma) palaeotemperatures as representing deeper burial followed by successive episodes of exhumation. for a palaeogeothermal gradient of 30°c/km and likely palaeo-surface temperatures, the late eocene palaeotemperatures require that the upper jurassic marine section in the borehole was buried below a 2750 m thick cover of upper jurassic – eocene rocks prior to the onset of late eocene exhumation. as these sediments are now near outcrop at c. 200 m above sea level, they have been uplifted by at least 3 km since maximum burial during post-rift thermal subsidence. the results are consistent with estimates of rock uplift on milne land since the late eocene and with interpretation of ocean drilling program (odp) data off south-east greenland suggesting that mid-cenozoic uplift of the margin triggered the marked influx of coarse clastic turbidites during the late oligocene above a middle eocene to upper oligocene hiatus. keywords: east greenland, jameson land, upper jurassic, apatite fission-track analysis, burial, exhumation ___________________________________________________________________________ p.f.g., geotrack international, 37 melville road, brunswick west, victoria 3055, australia. e-mail: mail@geotrack.com.au p.j., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark with sedimentary basins offshore east greenland yet to be drilled, the onshore jameson land basin (surlyk 2003) provides a window into the nature of potential jurassic source-rock sequences in offshore basins. understanding the thermal history and maturity development in the onshore sequences can therefore provide unique insights into the prospectivity of the offshore basins. the jameson land basin has itself been the focus of hydrocarbon exploration by a group of concessionaries with atlantic richfield (arco) as operator, but in 1990 the group decided not to continue exploration (christiansen et al. 1992; mathiesen et al. 2000). one of the main geological risks that arco critically assessed was the thermal maturity of their main target, the upper permian carbonates, since earlier studies had concluded that tertiary exhumation had removed up to 3 km of cretaceous sediments and paleocene–eocene volcanic rocks across the basin (christiansen et al. 1992). currently, greenland gas and oil a/s and nunaoil a/s hold a hydrocarbon exploration and exploitation licence across much of the jameson land basin (www.govmin.gl). presented here is a thermal history study of samples obtained from the core from the blokelv-1 borehole, drilled by geus in 2008 to provide detailed information about the upper jurassic stratigraphy and petroleum system of the jameson land basin (fig. 1). geus suc© geus, 2018. geological survey of denmark and greenland bulletin 42, 133–147. available at: www.geus.dk/bulletin42 mailto:mail@geotrack.com.au http://www.govmin.gl http://www.geus.dk/bulletin42 134134 cessfully drilled and cored the blokelv-1 borehole in the central part of the basin (ground level 200 m above sea level; a.s.l.), targeting the upper jurassic prolific sourcerock interval of the hareelv formation (bjerager et al. 2018a, this volume). the 231 m thick, marine succession cored in the blokelv-1 borehole covers the middle oxfordian to lower volgian interval. the core has 100% recovery, and it consists of interlayered organic-rich, laminated mudstones, massive sandstones and heterolithic sandstone–mudstone intervals of the katedralen member and massive sandstones of the sjællandselv member of the hareelv formation. the recovered core is of very high quality and has been subject to an extensive sampling and analytical programme designed to investigate aspects of the petroleum geology of the jameson land basin and published in nine papers (ineson & bojesenkoefoed (eds) 2018, this volume). the dataset provides an excellent reference for kimmeridge clay formationequivalent deposits in the north atlantic area. at first glance, the relatively low-lying landscape of the jameson land basin is in sharp contrast to the high terrains of the volcanic province along blosseville kyst to the south of scoresby sund, where elevations reach 3.7 km a.s.l. this might initially suggest that jameson fig. 1. a: simplified geology of the jameson land basin and adjacent areas showing the location of the blokelv-1 borehole (modified after surlyk 2003); inset shows the location of the study area in east greenland. bk: blosseville kyst. k: kangerlussuaq. b: stratigraphic column of the blokelv-1 borehole showing the stratigraphic level of the two samples used in this study (1, gc1052-1; 2, gc1052-2). log modified from bjerager et al. 2018a, this volume). litho: lithostratigraphic subdivision. chrono: chronostratigraphic subdivision. sj. mb: sjællandselv member. td: total depth. b c c' 74°n 76°n milne land profile in fig. 6 s c o r e s b y s u n d 25 km22°w24°w26°w 20°wliverpool land a b permian carboniferous pre-carboniferous fault sandstone mudstone igneous intrusion gc1052 sample blokelv-1 71°n 72°n palaeogene volcanics triassic cretaceous jurassicice k bk jameson land chrono.litho. volgian o xf or di an m id dl e up pe r lo w er up pe r k im m er id gia n u pp er ju ra ss ic h ar ee lv fo rm at io n ka te dr ale n m em be r sj. mb depth (m) 50 1 2 100 150 200 233.8 td pj 1 135 land has undergone a less complex post-jurassic history of uplift and erosion compared to regions to the south. however, thermochronological data from the jurassic sediments of the jameson land basin presented here (see also mathiesen et al. 2000; hansen et al. 2001) provide evidence of post-jurassic burial and exhumation that is surprisingly similar to the region around kangerlussuaq, to the south (larsen & saunders 1998; brooks 2011; bonow et al. 2014; japsen et al. 2014). thermal history interpretation afta data two sandstone core samples from the lower and upper levels of the blokelv-1 borehole (fig. 1b) were processed for apatite fission-track analysis (afta), and both samples gave excellent apatite yields. apatite fission track ages of 50.4 ± 6.2 ma and 38.5 ± 4.0 ma in the two samples are much less than the depositional age of the sampled units; at depths of less than 250 m (and present-day temperatures less than 20°c), this degree of age reduction immediately shows that the sampled units have been much hotter in the past. mean track lengths of 12.25 ± 0.22 µm and 11.71 ± 0.23 µm also demonstrate that these samples have been hotter in the past, prior to cooling to present-day temperatures. full details of the afta data are provided in appendix 1. quantitative thermal history constraints have been extracted from the data using principles outlined by green & duddy (2012) and green et al. (2013), with results summarised in table 1. afta data from both samples provided highly reliable thermal history constraints. thermal history interpretation of afta data the afta data in sample gc1052-1 can be explained in terms of two palaeo-thermal episodes, as detailed in table 1. in contrast, the afta data in sample gc10522 require three palaeo-thermal episodes to explain all aspects of the data, although the precise timing of the earliest episode cannot be defined with confidence. this is because the palaeotemperature of 100–110°c in the second episode produced almost total annealing of all tracks formed up to that time, largely masking the previous history. on the basis of evidence discussed below, we table 1. palaeotemperature analysis summary: afta and vr data from the blokelv-1 borehole sample mean present stratigraphic vr** maximum onset maximum onset number depth temperature* age (%) palaeotemp.+ of cooling+ palaeotemp.+ of cooling+ gc1052(m below kb) (°c) (ma) (°c) (ma) (°c) (ma) 1 6 5 159–146 100–105 58–28 70–80 17–5 23.59 159–146 0.52 86 32.78 159–146 0.54 90 68.77 159–146 0.50 83 80.77 159–146 0.55 91 92.74 159–146 0.56 93 104.79 159–146 116.82 159–146 0.56 96 152.75 159–146 0.61 100 164.82 159–146 0.68 113 176.77 159–146 0.62 102 188.77 159–146 0.65 108 200.76 159–146 0.66 109 212.77 159–146 0.78 125 2 219 12 159–146 >110 55–50? 100–110 50–28 80–85 13–3 224.74 159–146 0.74 121 combined timing (ma): 55–50 13–5 50–28 * present temperature estimates based on an assumed surface temperature of 4°c, and an assumed thermal gradient of 30°c/km. ** from bojesen-koefoed et al. (2018). + thermal history interpretation of afta data is based on an assumed heating rate of 1°c/myr and a cooling rate of 10°c/myr. quoted ranges for palaeotemperature and onset of cooling correspond to ±95% confidence limits. conditions shown in italics represent events that cannot be rigorously defined from the afta data. pj table 1 136136 infer that the earlier event in which this sample cooled below 110°c was related to igneous activity in the region (larsen 2018, this volume), in the interval 55 to 50 ma. timing constraints derived from afta data in each sample are listed in table 1, and in table 2 these constraints are compared with the timing of three cenozoic cooling episodes defined from afta data in two previous studies of the east greenland margin: (a) a study of the region north of the jameson land basin (thomson et al. 1999) and (b) a study in the kangerlussuaq region to the south of jameson land ( japsen et al. 2014). the similarity in timing of the late eocene and the late miocene cooling episodes defined in this study and in the previous studies suggests that each represents a regional, synchronous cooling episode across the entire region, and by combining all constraints we arrive at our preferred timing of the onset of cooling in these events, between 40 and 35 ma and ~10 ma (table 2). in addition to the episodes shown in fig. 2, japsen et al. (2014) also defined late oligocene and early miocene palaeo-thermal episodes in the region around kangerlussuaq, but these are restricted to that region and do not extend to the jameson land basin, so are not considered here. we interpret the three events illustrated in fig. 2 (from table 2) in the following way: 55–50 ma event. the timing of the early eocene (55–50 ma) event overlaps with that of intensive palaeogene intrusive activity and correlates with the age of c. 53 ma for dykes and sills in the jameson land basin (hald & tegner 2000; larsen 2018, this volume). palaeotemperatures associated with this event are 100°c or above and are identified sporadically around the region ( japsen et al. 2014). on this basis, the palaeotemperatures characterising this episode are interpreted to be due either to contact or hydrothermal effects associated with igneous activity. no convincing evidence for any regional paleocene to mid-eocene cooling (related to exhumation) has been identified for samples in the area around jameson land (thomson et al. 1999; japsen et al. 2014), and the geological history recorded south of jameson land indicates that subsidence and burial dominated at the palaeocene–eocene transition (brooks 2011; bonow et al. 2014). 40–35 ma event. late eocene cooling beginning between 40 and 35 ma was interpreted largely in terms of regional uplift resulting in kilometre-scale exhumation by both japsen et al. (2014) in the kangerlussuaq region and thomson et al. (1999) in the region to the north of jameson land. japsen et al. (2014) interpreted the endresult of this phase of exhumation to have been a regional peneplain, the upper planation surface (ups) of bonow et al. (2014). the presence of late eocene intrusive bodies around traill ø (price et al. 1997) suggests the possibility locally of a significantly elevated basal heat flow in this region at this time. 10 ma event. late miocene cooling beginning at c. 10 ma was again detected by both japsen et al. (2014) in afta data 55–50 50–28 13–5 blokelv samples gc1052-1, -2 afta data (study a) 40–30 10–5 afta data (study b) * 40–35 ~10 age of intrusion (study c) ~53** (55–51) preferred regional timing 55–50 40–35 ~10 (early eocene) (late eocene) (late miocene) dominant mechanism of cooling cooling after intrusive heating exhumation exhumation table 2. intervals defining the beginning of episodes of cooling from afta data onset of cooling (ma) study a. 11 outcrop samples from northern east greenland, north of the jameson land basin, 72–74°n (thomson et al. 1999). study b. 90 samples from outcrops and drillholes in a regional study of southern east greenland focussed between 68 and 70°n (japsen et al. 2014). study c. analysis of the igneous intrusions in the blokelv cored borehole (larsen 2018, this volume). *japsen et al. (2014) related an event of cooling with overlapping timing (55–50 ma) in southern east greenland to the emplacement of the kangerlussuaq intrusion. **the blokelv sills are tholeiitic basalts considered to belong to the main group of dykes and sills in the jameson land basin which has been dated at ~53 ma. the intrusions form part of a 55–51 ma group of tholeiitic basalt intrusions that were emplaced within the sedimentary basins in east greenland (larsen 2018, this volume). pj table 2 137 the kangerlussuaq region and thomson et al. (1999) to the north of jameson land. in both cases cooling was interpreted in terms of regional uplift resulting in kilometre-scale exhumation. outcrop samples at low elevations around kangerlussuaq cooled from peak palaeotemperatures around 60–70°c at this time. japsen et al. (2014) interpreted these values to represent burial below the ups that defines the present-day surface of elevated summits along the blosseville kyst. table 2 illustrates a high degree of consistency between the timing of cooling identified in this study and the dominant regional episodes identified by thomson et al. (1999) and japsen et al. (2014). on this basis, the results from sample gc1052-1 are interpreted as representing the two most recent episodes (i.e. late eocene and late miocene), while these two episodes as well as the early eocene episode are recognised in sample gc1052-2. thermal history interpretation of vitrinite reflectance data results of vitrinite reflectance (vr) analyses are plotted against depth below surface in fig. 3, together with equivalent vr (vreq) values derived from rock-eval tmax values (bojesen-koefoed et al. 2018, this volume). also shown are the ranges of equivalent vr values (vreq) derived from afta data in samples gc1052-1 and -2 (defined by the maximum paleotemperature in each sample), together with the vr profile predicted on the basis 0 0255075100125 time (ma) 150175200225250 50 100 150 200 250 d ep th (m ) q gc1052-1 stratigraphic age regional cooling episodes: regional constraints on the onset of cooling palaeothermal constraints from afta sample triassic jurassic cretaceous cenozoic gc1052-2 55–50 ma 40–35 ma ~10 ma pj 2 fig. 2. timing constraints on cooling episodes derived from afta data in two samples from the blokelv-1 borehole (horizontal bars) plotted against depth. the regional constraints on the onset of cooling in three palaeo-thermal episodes (vertical bars) are based on afta data in the blokelv samples and in regional studies (table 2). q: quaternary. fig. 3. maturity indicators in samples from the blokelv-1 borehole plotted against depth; vr values (table 1) together with equivalent vr values (vreq) derived from rock-eval tmax data and equivalent ranges of vr defined from the afta data. note that the orange arrow for the deeper sample indicates that the constraints on maturity from the afta data only provide a lower limit. the solid sub-vertical line shows the profile predicted from the “default thermal history”, i.e., the history calculated from the assumption that all units throughout the well are currently at their maximum temperatures since deposition. the horizontal lines indicate the position of thin basaltic sills intruded into the hareelv formation (note the maturity effects close to these intrusions). 0 20 40 60 80 120 140 160 180 200 220 240 100 0.2 0.3 0.4 0.5 0.7 1.0 2.0 gc1052-2 gc1052-1 d ep th (m ) default history maturity (%ro) u pp er ju ra ss ic sill sill sill measured vr values (bojesen-koefoed et al. 2018, this volume) equivalent vr from afta equivalent vr from tmax afta sample horizon pj 3 138138 of the default thermal history (i.e. the history expected if the section has never been any hotter than it is today). both vr and vreq data plot well above the profile predicted by the default thermal history, confirming the evidence from afta that the sampled units have been hotter than their present-day temperatures at some time since deposition. mean vr values tend to be slightly lower than the vreq values derived from the tmax data throughout the section. in general, tmax values tend to be sensitive to a range of factors and are not used quantitatively to provide estimates of maximum palaeotemperature in the way that vr data are. note that the tmax data define local contact effects due to the three recognised intrusions, which the vr values do not show. maximum palaeotemperatures derived from the vr values (based on burnham & sweeney 1989) are listed in table 2, and show a progressive downhole increase from 86°c to 125°c through the section intersected in the borehole. this apparent increase of c. 40°c over a depth interval of c. 200 m is equivalent to a thermal gradient of 200°c/km, which is well outside the range of typical sedimentary basin thermal gradients (allen & allen 2013). integration of results from afta with the vr data resolves this apparent anomaly, as explained below. integration of afta and vr data, palaeotemperature profiles and mechanisms of heating and cooling palaeotemperature constraints from afta and the measured vr values are plotted against depth in fig. 4. maximum palaeotemperatures derived from vr values at depths between 20 m and 120 m are c. 10–20°c lower than the late eocene palaeotemperature range of 100–105°c indicated by afta data in sample gc10521 at similar depths. in contrast, maximum palaeotemperatures of c. 125°c derived from the two deepest vr values at depths of 210–225 m are higher than the late eocene palaeotemperatures derived from afta in sample gc1052-2 at a similar depth. they are, however, consistent with the early eocene palaeotemperature defined from afta data in this sample. furthermore, as illustrated in fig. 4, maximum palaeotemperatures derived from vr values between 150 m and 200 m are consistent with the trend of the late eocene palaeotemperatures derived from afta in samples gc1052-1 and -2. we interpret the mismatch at shallow depths between vr values and the late eocene palaeotemperature range from afta in sample gc1052-1 as due to suppression of reflectance levels at these depths. suppression is commonly observed in upper jurassic organic-rich mudstones of the north atlantic region (wilkins et al. 1992; newman 1997). bjerager et al. (2018b, this volume) and bojesen-koefoed et al. (2018, this volume) provided detailed discussions of the differences between the organic material above and below c. 100 m in the cored interval. they showed that at shallow depths, amorphous marine 0 20 40 60 80 100 120 140 160 180 200 220 240 260 temperature (°c) 0 50 100 150 200 250 d ep th (m ) afta sample horizon (with sample number) maximum palaeotemperature from vr estimates of maximum post-depositional palaeotemperatures from afta sill sill sill u pp er ju ra ss ic gc1052-2 gc1052-1 local heating? 30°c/km30°c/km low due to suppression of vr? thermal gradient (°c/km) 100 80 60 40 20 pj 4 fig. 4. interpreted palaeotemperature profiles describing the palaeotemperatures in two episodes derived from afta and vr data in the blokelv-1 borehole. based on evidence from regional studies of east greenland (table 2), we interpret the early eocene (55–50 ma) maximum palaeotemperature (green arrow) revealed by afta data in sample gc1052-2 to represent localised heating due to intrusive activity. the late eocene (40–35 ma) and late miocene (c. 10 ma) palaeotemperatures (red and blue horizontal bars, respectively) revealed by afta are interpreted to represent the effects of burial, with late eocene cooling representing the onset of regional, post-jurassic exhumation, and late miocene cooling representing the final phase of exhumation. it is thus inferred that the vr values recorded above 120 m in the borehole (yellow datapoints) are anomalously low due to suppression of reflectance in the sourcerock facies (organic-rich mudstones) (wilkins et al. 1992; newman 1997). we regard the vr values below 120 m (red datapoints) as reliable indications of the degree of post-depositional heating. linear profiles (red and blue lines) represent our preferred interpretation (based on regional data) involving palaeogeothermal gradients of ~30°c for the late eocene and miocene episodes. 139 organic material dominate whereas an increased content of terrigenous organic material occurs at depth. as illustrated in fig. 4, maximum palaeotemperatures from vr data at depths between c. 150 m and 200 m, together with late eocene palaeotemperatures derived from afta in samples gc1052-1 and -2 define a linear profile characterised by a palaeogeothermal gradient around 30°c/km. late miocene palaeotemperatures defined from afta in the two samples are also consistent with a similar palaeogeothermal gradient. this gradient contrasts markedly with the apparent gradient of c. 200°c/ km defined by the vr data (see above), and is more typical of heating related to deeper burial. accepting this interpretation, units throughout the borehole underwent a major phase of cooling in the late eocene (beginning between 40 and 35 ma), followed by a later phase of cooling in the late miocene (c. 10 ma). most units cooled from maximum post-depositional palaeotemperatures in the late eocene event but locally some horizons reached higher palaeotemperatures in the early eocene, presumably reflecting the effects of intrusive bodies. this interpretation is consistent with the results reported from the kangerlussuaq region to the south by japsen et al. (2014), who regarded the late miocene and late eocene cooling episodes as representing successive episodes of exhumation, while early eocene events were interpreted as local effects associated with igneous intrusions. thermal history synthesis on the basis of the discussion presented above, we interpret the palaeotemperature constraints derived from afta and vr data in the blokelv-1 borehole as representing the combined effects of deeper burial followed by successive episodes of exhumation in the late eocene and miocene, as well as localised early eocene heating due either to contact heating or hydrothermal effects associated with intrusive activity. the results provided here are highly consistent with regional data and the interpretation presented here is regarded as reliable. thermal history reconstruction here we present reconstructed thermal and burial-uplift histories for the upper jurassic section intersected in the blokelv-1 borehole (fig. 5) based on the results presented above. it should be emphasised that while the preferred reconstruction illustrated here provides a satisfactory explanation of the afta and vr data from this well, the reconstruction is not unique. therefore, it is important to appreciate those aspects of the histories that are constrained by the data, and those that are not. factors that can be confidently defined in this study (within the limits of analytical uncertainty) include: (a) magnitude of heating at the palaeothermal maximum and the subsequent palaeo thermal peaks and (b) timing of the onset of cooling in each episode. factors that can be defined when samples are available over a sufficiently large range of depths or elevations, but cannot be constrained in this study include: (a) palaeogeothermal gradients during each episode and (b) additional burial during each episode for a specified value of palaeogeothermal gradient. aspects which cannot be uniquely defined in any situation include: (a) thermal history prior to the palaeo-thermal maximum and/or the subsequent palaeo-thermal peak, (b) amounts of re-burial between multiple episodes within a single unconformity and (c) detailed style of cooling history from each episode. figure 5a illustrates a possible thermal history reconstruction for the sedimentary units intersected in the blokelv-1 borehole, based on the synthesis developed above. key aspects of this reconstruction are: • surface temperature of 20°c at 30 ma and earlier, decreasing to 10°c at 10 ma and to a present-day value of 4°c over the last 12 myr. • palaeogeothermal gradient of 30°c/km, constant to the present day. • localised heating within the vicinity of sample gc1052-2 to a palaeotemperature around 120°c, shown at c. 53 ma but any time between 56 and 45 ma is allowed by the regional timing constraints on this episode. • an additional 2750 m of section deposited above the upper jurassic section intersected in the borehole, between 146 and 35 ma (resulting in heating of sample gc1052-1 to c. 100°c prior to late eocene cooling and exhumation). • subsequent removal of 1150 m (arbitrary) of section between 35 and 30 ma, followed by deposition of a further 600 m (arbitrary) of section between 30 and 10 ma. • removal of the remaining 2200 m of additional section between 10 ma and the present day. note that the amount of re-burial between the two episodes of exhumation cannot be controlled by the data, 140140 and therefore also the amount of section removed in the initial episode beginning at 35 ma is also not constrained. this reconstruction is considered to provide a reliable depiction of the post-depositional history of the upper jurassic section intersected in the blokelv-1 borehole. note that in this reconstruction, for the purposes of illustration, the localised heating around sample gc1052-2 at c. 53 ma is shown as taking place over a duration of 2 myr, as is the subsequent cooling. however, in reality heating and cooling would have been much more rapid, and for that reason the true maximum palaeotemperatures would have been much higher. comparison with previous studies magnitude of exhumation christiansen et al. (1992) presented the results of a first study of the exhumation of the jameson land basin. they based their estimate on a range of observations including thermal maturity parameters, apatite fissiontrack parameters, porosity and seismic velocities; in particular they reported vr values up to 0.6% for organicrich permian and jurassic formations at outcrop. they interpreted their data to indicate that between 1.5 and 3 km of overburden had been removed across the basin, 0 20 40 60 80 100 160 140 120 100 80 time (ma) 60 40 20 0 120 te m pe ra tu re (º c ) possible reconstruction j k pa pe o m gc1052-1 gc1052-2 ? possible reconstruction j k pa pe o m 0 2000 1500 1000 500 2500 3000 bu ria l d ep th (m ) 160 140 120 100 80 time (ma) 60 40 20 0 regional constraints on the onset of cooling palaeothermal constraints from afta in sample gc1052-1 palaeothermal constraints from afta in sample gc1052-2 thermal histories for the two sample horizons approximate burial depths during late eocene and late miocene palaeothermal peaks ? a b pj 5 fig. 5. schematic illustration of the preferred thermal (a) and burial (b) history reconstruction for the section intersected in the blokelv-1 borehole based on the afta and vr data. j: jurassic. k: cretaceous. pa: palaeocene. e: eocene. o: oligocene. m: miocene. p: pliocene. a: comparison of palaeothermal constraints from afta in two samples (red and blue boxes) with thermal histories of the corresponding sample horizons (red and blue curves). the vertical columns show the preferred timing of three dominant palaeothermal episodes identified in the region (fig. 2; table 2). black line at the top of the diagram defines the assumed palaeosurface temperature. b: burial and exhumation history corresponding to a; the blue boxes show the approximate depth of burial during the late eocene and late miocene palaeothermal peaks. note that no convincing evidence for regional paleocene to mid-eocene exhumation has been identified for samples in the area around jameson land (thomson et al. 1999; japsen et al. 2014). 141 and that the lost cover likely consisted of more than 1 km of cretaceous sediments as well as 1–2 km of palaeogene basalts (primarily across the southern part of the basin). mathiesen et al. (2000) investigated the denudation history of the jameson land basin using basin modelling constrained by apatite fission-track data. they concluded that the upper jurassic sediments that are exposed at the present day across jameson land were buried below a 2–3 km thick rock column: (1) a cretaceous succession that varied from 1.3 km in the south to 0.3 km in the north; (2) a wedge of palaeogene volcanics with a thickness of >2 km in the south thinning to <0.1 km in the north. according to their calculations, the erosion happened in response to tectonic uplift of c. 1 km. the magnitude of the section removed at the location of the blokelv-1 borehole defined in this study (c. 2.8 km) thus agrees well with that estimated by mathiesen et al. (2000) (removal of 2 –3 km of section across jameson land). hansen et al. (2001) studied the late mesozoic – cenozoic thermal history of the jameson land basin constrained by apatite and zircon fission-track data of permian to jurassic sedimentary rocks at outcrop. these authors interpreted their results in terms of regional thermal evolution related to burial leading to temperatures close to and in excess of the maximum temperatures of the apatite annealing interval (c. 125°c) followed by cooling mainly due to cenozoic uplift and erosion. furthermore, basaltic dyke and sill intrusions were found locally to cause resetting of apatite fission-track ages. these results are thus broadly in agreement with the conclusions presented here. our estimate of the magnitude of the section removed above the blokelv-1 borehole (c. 2.8 km) agrees well with the results of bonow et al. (2014) and japsen et al. (2014) in their studies of the area between milne land and kangerlussuaq (68–71°n) based on integration of evidence from stratigraphic landscape analysis, thermochronology and the stratigraphic record (fig. 6). these authors argued that the present-day high elevation in east greenland is the result of three tectonic phases of uplift and erosion during the cenozoic that followed the eruption of voluminous flood basalts onto a largely horizontal lava plain near sea level at the paleocene– eocene transition (larsen & saunders 1998; brooks 2011; bonow et al. 2014), viz: 1. the late eocene (c. 35 ma) phase of uplift and erosion led to formation of an oligo–miocene erosion surface (peneplain) near sea level, the upper planation surface (ups). 2. uplift of this surface in the late miocene (c. 10 ma) led to formation of a lower surface (the lower planation surface, lps) by incision below the uplifted ups. 3. an early pliocene uplift phase (c. 5 ma) led to incision of valleys and fjords below the lps, resulting in mountain peaks reaching 3.7 km a.s.l. today, remnants of the ups are preserved west of jameson land near the summits of milne land at c. 2 km a.s.l. bonow et al. (2014) estimated the magnitude of rock uplift at milne land to be c. 2.7 km based on larsen et al.’s (1989) investigation of zeolite isograds (levels of equal thermal alteration; see also neuhoff et al. 1997), and on their argument that the absence of the shallow and less altered zeolite zones may be explained by the removal of these zones by erosion. for milne land, they concluded that a section of about 900 m had been removed above the basalt flows that cover the summits there (assuming a palaeogeothermal gradient of 40°c at the time of zeolite formation). bonow et al. (2014) then assumed that the palaeo-surface during the formation of the zeolites was near sea level (shortly after the eruption of the volcanics 3 milne land land surface prior to onset of late eocene denudation relative to present-day sea level removed section present-day land area w e jameson land ? liverpool land 25 km blokelv borehole 2 el ev at io n (k m a .s. l.) 1 0 pj 6 fig. 6. present-day elevation profile with indication of the section removed since late eocene maximum burial. based on bonow et al. (2014) and japsen et al. (2014). for location of profile, see fig. 1. 142142 at the paleocene–eocene transition at a time of regional subsidence; brooks 2011). consequently, the basalts, now at c. 1800 m a.s.l., have been uplifted 2700 m (i.e. 1800 + 900 m) since the formation of the zeolites in the eocene. japsen et al. (2014) found that the ups that defines the summits of milne land at c. 2 km a.s.l. represents the remnants of the oligo-miocene peneplain that was formed by erosion to sea level after late eocene uplift and erosion, and that this peneplain was uplifted to its present elevation during uplift that began in the late miocene. assuming that the lost cover of 900 m (estimated from zeolite isograds) had been removed after late eocene maximum burial, the magnitude of rock uplift was about 2.9 km since the late eocene. these estimates of the magnitude of rock uplift thus match those presented here because the upper jurassic marine sediments at outcrop (now at c. 200 m a.s.l.) at the blokelv-1 location have been uplifted by a minimum of 3.0 km since their maximum burial below a (now lost) cover of c. 2.8 km in the late eocene. timing of burial and exhumation with no way of determining the timing of the onset of exhumation, christiansen et al. (1992) and mathiesen et al. (2000) assumed that the exhumation of the jameson land basin accelerated after the palaeogene volcanic eruptions. the present study, however, clearly shows that the onset of exhumation was in the late eocene (c. 35 ma) and documents that the exhumation process took place in at least two stages. these results thus imply that the section of rocks removed across jameson land not only included volcanic rocks extruded at the paleocene– eocene transition (as assumed by mathiesen et al. 2000) but also a sedimentary cover deposited during the 20 myr that followed the volcanic eruptions till the onset of exhumation in the late eocene. the geological record from the area south of jameson land confirms the timing of the cenozoic events of burial and exhumation presented here. the area between scoresby sund and kangerlussuaq subsided after continental breakup at 56 ma. this is documented by the igtertivâ formation which immediately overlies the main basalts of larsen et al. (1989, 2013) and includes volcanic flows interdigitated with marine sediments. the volcanic pile of the main basalts is up to 6 km thick and was erupted in less than 1 myr (pedersen et al. 1997). subsidence continued during deposition of the fluvial to shallow-marine kap dalton group (early to mid lutetian), which interfingers with and overlies the igtertivâ formation in a downfaulted block at kap dalton (larsen et al. 1989, 2005, 2013). results from odp site 918, off se greenland (leg 152, c. 63°n) show that major uplift of the margin occurred long after continental break-up. larsen et al. (1994) reported that the marine, lower eocene sediments drilled there indicated low sedimentation rates with limited terrigenous influx before a middle eocene to upper oligocene hiatus. larsen et al. (1994) thus argued that mid-cenozoic uplift of the inner margin triggered the sudden, voluminous influx of coarse clastic turbidites at this odp site during the late oligocene. uplift phases in the late miocene and in the pliocene are consistent with the late miocene, early pliocene and middle pliocene ages of seismic sequence boundaries within the late neogene and quaternary deep-sea sedimentary succession off se greenland (clausen 1998). conclusions three palaeo-thermal episodes affected the upper jurassic sediments penetrated by the blokelv-1 borehole based on afta and vr data combined with regional afta data. these episodes are interpreted to be due to the following mechanisms: • early eocene (55–50 ma) palaeotemperatures represent localised early eocene heating related to intrusive activity. • late eocene (40–35 ma) palaeotemperatures represent deeper burial followed by exhumation. • late miocene (c. 10 ma) palaeotemperatures represent deeper burial followed by exhumation. the presence of two elevated planation surfaces in the region that were formed and uplifted after the volcanic eruptions supports the interpretation of the palaeothermal data in terms of episodic rather than monotonic cooling (bonow et al. 2014; japsen et al. 2014). the late eocene palaeotemperatures require that c. 2800 m of upper jurassic – eocene rocks covered the upper jurassic section in the borehole prior to the onset of late eocene exhumation, assuming a palaeogeothermal gradient of 30°c/km and likely palaeo-surface temperatures. this implies that maximum burial in the jameson land basin was achieved long after the rift climax in central east greenland at the jurassic–cretaceous tran143 sition (surlyk 2003) and after the volcanic eruptions that accompanied break-up in the north-east atlantic at the paleocene–eocene transition (pedersen et al. 1997). as the upper jurassic sediments at the location of the blokelv-1 borehole now crop out at c. 200 m a.s.l., they have been uplifted by at least 3 km since maximum burial during post-rift thermal subsidence. such a magnitude of rock uplift is comparable with estimates from milne land where a regional peneplain (the upper planation surface, ups) defines the summits at c. 2 km a.s.l. this surface was formed by erosion to sea level (after removal of a rock column of c. 900 m) following late eocene uplift and subsequently uplifted to its present elevation during uplift that began in the late miocene. consequently, the rock uplift on milne land was c. 2.9 km since the late eocene. rock uplift in the order of 3 km since the late eocene has thus affected a wide area, far beyond the boundaries of the jameson land basin. that strong uplift of the east greenland margin began at the eocene–oligocene transition is supported by interpretation of odp data off south-east greenland which suggest that uplift of the margin at this time triggered the marked influx of coarse clastic turbidites during the late oligocene above a middle eocene to upper oligocene hiatus. acknowledgements we acknowledge the pertinent and constructive comments of the referees, andrew carter and andrew g. whitham. jette halskov and stefan sølberg are thanked for graphical support. reference list allen, p.a. & allen, j.r. 2013: basin analysis: principles and application to petroleum play assessment, 632 pp. indianapolis: john wiley & sons. bjerager, m., bojesen-koefoed, j. & piasecki, s. 2018a: the upper jurassic blokelv-1 cored borehole in jameson land, east greenland – an introduction. in: ineson, j. & bojesen-koefoed, j.a. (eds): petroleum geology of the upper jurassic – lower cretaceous of east and north-east greenland: blokelv-1 borehole, jameson land basin. geological survey of denmark and greenland bulletin 42, 7–14 (this volume). bjerager, m., kjøller, c., olivarius, m., olsen, d. & schovsbo, n. 2018a: sedimentology, geochemistry and reservoir properties of upper jurassic deep marine sediments (hareelv formation) in the blokelv-1 borehole, jameson land basin, east greenland. in: ineson, j. & bojesen-koefoed, j.a. (eds): petroleum geology of the upper jurassic – lower cretaceous of east and north-east greenland: blokelv-1 borehole, jameson land basin. geological survey of denmark and greenland bulletin 42, 39–64 (this volume). bojesen-koefoed, j.a., peter nytoft, h.p., petersen, h.i., piasecki, s. & pilgaard, a. 2018: petroleum potential of the upper jurassic hareelv formation, jameson land, east greenland. in: ineson, j. & bojesen-koefoed, j.a. 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(eds): petroleum geology of the upper jurassic – lower cretaceous of east and north-east greenland: blokelv-1 borehole, jameson land basin. geological survey of denmark and greenland bulletin 42, 168 pp. (this volume). japsen, p., green, p.f., bonow, j.m., nielsen, t.f.d., & chalmers, j.a. 2014: from volcanic plains to glaciated peaks: burial and exhumation history of southern east greenland after opening of the ne atlantic. global and planetary change 116, 91–114. larsen, h.c. & saunders, a.d. 1998: tectonism and volcanism at 144144 the southeast greenland rifted margin: a record of plume impact and later continental rupture. in: saunders, a.d. et al. (eds): proceedings of the ocean drilling program, scientific results 152, 503–533. larsen, h.c., saunders, a.d., clift, p.d. et al. 1994: 13. summary and principal results, proceedings of the ocean drilling program, initial reports 152, 279–292. larsen l.m. 2018: igneous intrusions in the cored upper jurassic succession of the blokelv-1 core, jameson land basin, east greenland. in: ineson, j & bojesen-koefoed, j.a. (eds): petroleum geology of the upper jurassic – lower cretaceous of east and north-east greenland: blokelv-1 borehole, jameson land basin. geological survey of denmark and greenland 42, 127–132 (this volume). larsen, l.m., pedersen, a.k., sørensen, e.v., watt, w.s. & duncan, r.a. 2013: stratigraphy and age of the eocene igtertivâ formation basalts, alkaline pebbles and sediments of the kap dalton group in the graben at kap dalton, east greenland. geological society of denmark bulletin 61, 1–18. larsen, l.m., watt, w.s. & watt, m. 1989: geology and petrology of the lower tertiary plateau basalts of the scoresby sund region, east greenland. grønlands geologiske undersøgelse bulletin 169. 164 pp. larsen, m., heilmann-clausen, c., piasecki, s. & stemmerik, l. 2005: at the edge of a new ocean: post-volcanic evolution of the palaeogene kap dalton group, east greenland. in: doré, a.g. & vining, b. (eds): petroleum geology: north-west europe and global perspectives-proceedings of the 6th petroleum geology conference. geological society (london) 923–932. mathiesen, a., bidstrup, t. & christiansen, f.g. 2000: denudation and uplift history of the jameson land basin, east greenland – constrained from maturity and apatite fission data. global and planetary change 24, 275–301. neuhoff, p.s., watt, w.s., bird, d.k. & pedersen, a.k. 1997: timing and structural relations of regional zeolite zones in basalts of the east greenland continental margin. geology 25, 803–806. newman, j. 1997: new approaches to detection and correction of suppressed vitrinite reflectance. australian petroleum production & exploration association (appea) journal 27, 524–535. pedersen, a.k., watt, m., watt, w.s. & larsen, l.m. 1997: structure and stratigraphy of the early tertiary basalts of the blosseville kyst, east greenland. journal of the geological society (london) 154, 565–570. price, s., brodie, j., whitham, a. & kent, r. 1997: mid-tertiary rifting and magmatism in the traill o region, east greenland. journal of the geological society (london) 154, 419–434. surlyk, f. 2003: the jurassic of east greenland: a sedimentary record of thermal subsidence, onset and culmination of rifting. in: ineson, j. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 659–722. thomson, k., green, p.f., whitham, a.g., price, s.p. & underhill, j.r. 1999: new constraints on the thermal history of north-east greenland from apatite fission-track analysis. geological society of america bulletin 111, 1054–1068. wilkins, r.w.t., wilmshurt, j.r., russel, n.j., hladky, g., ellacott, m.w. & buckingham, c. 1992: fluorescence alteration and the suppression of vitrinite reflectance. organic geochemistry 18, 629–640. _________________________________________________________________________________________ manuscript received 16 december 2015; revision accepted 8 february 2018 145 slide current ns ni na ρs ρi ratio u cl ft age ref grain no (ppm) (wt%) (ma) g1124-9 3 1 23 36 4.414e+04 1.015e+06 0.043 8.1 0.04 12.2 ± 12.4 g1124-9 6 14 51 24 9.270e+05 3.377e+06 0.275 27.0 0.03 76.4 ± 23.1 g1124-9 7 0 13 49 0.000e+00 4.216e+05 0.000 3.4 0.14 0.0 ± 36.1 g1124-9 10 29 129 30 1.536e+06 6.833e+06 0.225 54.7 0.01 62.6 ± 13.0 g1124-9 11 18 48 24 1.192e+06 3.178e+06 0.375 25.4 0.06 104.1 ± 28.9 g1124-9 12 6 35 60 1.589e+05 9.270e+05 0.171 7.4 0.05 47.8 ± 21.2 g1124-9 13 27 96 60 7.151e+05 2.543e+06 0.281 20.3 0.02 78.2 ± 17.2 g1124-9 14 6 53 42 2.270e+05 2.005e+06 0.113 16.0 0.03 31.6 ± 13.6 g1124-9 15 54 339 20 4.290e+06 2.693e+07 0.159 215.6 0.11 44.4 ± 6.6 g1124-9 16 4 33 35 1.816e+05 1.498e+06 0.121 12.0 0.00 33.8 ± 17.9 g1124-9 17 1 13 60 2.648e+04 3.443e+05 0.077 2.8 0.08 21.5 ± 22.3 g1124-9 18 9 110 42 3.405e+05 4.162e+06 0.082 33.3 0.03 22.9 ± 8.0 g1124-9 19 5 18 30 2.648e+05 9.534e+05 0.278 7.6 0.02 77.3 ± 39.1 g1124-9 20 19 49 80 3.774e+05 9.733e+05 0.388 7.8 0.02 107.6 ± 29.2 g1124-9 21 4 32 70 9.080e+04 7.264e+05 0.125 5.8 0.02 34.9 ± 18.5 g1124-9 22 31 234 80 6.158e+05 4.648e+06 0.132 37.2 0.03 37.0 ± 7.1 g1124-9 23 11 113 50 3.496e+05 3.591e+06 0.097 28.7 0.03 27.2 ± 8.6 g1124-9 25 4 27 50 1.271e+05 8.581e+05 0.148 6.9 0.02 41.3 ± 22.2 g1124-9 28 1 6 100 1.589e+04 9.534e+04 0.167 0.8 0.07 46.5 ± 50.2 g1124-9 32 25 93 50 7.945e+05 2.956e+06 0.269 23.7 0.21 74.8 ± 17.0 269 1515 4.309e+05 2.427e+06 19.4 for abbreviations, see page 147 area of basic unit = 6.293e+07 cm2 χ2 = 47.992 with 19 degrees of freedom p(χ2) = 0.0% age calculated using a zeta of 392.9 ± 7.4 for cn5 glass ρd = 1.424e+06 cm2 nd = 2241 ρd interpolated between top of can; ρd = 1.388e+06 cm2 nd = 1092 ρd interpolated between bottom of can; ρd = 1.461e+06 cm2 nd = 1149 age dispersion = 38.770% ns / ni = 0.178 ± 0.012 mean ratio = 0.176 ± 0.024 pooled age = 49.5 ± 3.6 ma central age = 50.4 ± 6.2 ma 0 50 100 150 0.0 0.1 0.2 0 -2 +2 76 0 43 21 1.00 0.75 0.50 0.25 0.00 0.0 0.5 1.0 1.5 2.0 wt% cl wt% cl 30 20 10 0 0 5 10 15 20 track length (µm) n fr ac tio n fi ss io ntr ac k ag e (m a) ml: 12.25 µm std dev: 2.28 ft age: 50.4 ma stratigraphic age 107 200 0.25 radial plot of single grain ages* distribution of chlorine (cl) contents in apatite grains single grain age vs. weight % cl for individual apatite grains distribution of confined track lengths pj appendix sample 1 * see appendix b in geotrack report gc1052 (available online) for details of radial plot construction. colour datapoints indicate wt% cl: dark blue: <0.1%. green: 0.1-0.2%. pale blue: 0.2-0.3%. appendix 1: analytical details of the afta data sample gc1052-1: apatite this appendix documents the raw fission-track count data together with radial/compositional graphical plots. the location of the two samples (gc1052-1, gc1052-2) in the blokelv-1 core is indicated in fig. 1. 146146 slide current ns ni na ρs ρi ratio u cl ft age ref grain no (ppm) (wt%) (ma) g1124-10 3 15 84 50 4.767e+05 2.670e+06 0.179 21.3 0.23 49.9 ± 14.1 g1124-10 4 5 31 70 1.135e+05 7.037e+05 0.161 5.6 0.20 45.1 ± 21.8 g1124-10 10 7 56 50 2.225e+05 1.780e+06 0.125 14.2 0.01 35.0 ± 14.1 g1124-10 11 9 126 30 4.767e+05 6.674e+06 0.071 53.2 0.01 20.0 ± 6.9 g1124-10 12 3 63 28 1.703e+05 3.575e+06 0.048 28.5 0.02 13.4 ± 7.9 g1124-10 14 41 198 35 1.861e+06 8.990e+06 0.207 71.7 0.08 57.9 ± 10.1 g1124-10 15 14 94 36 6.180e+05 4.149e+06 0.149 33.1 0.01 41.7 ± 12.0 g1124-10 16 1 12 28 5.675e+04 6.810e+05 0.083 5.4 0.02 23.4 ± 24.3 g1124-10 17 9 48 48 2.980e+05 1.589e+06 0.188 12.7 0.10 52.4 ± 19.1 g1124-10 19 5 54 30 2.648e+05 2.860e+06 0.093 22.8 0.03 25.9 ± 12.1 g1124-10 21 16 152 40 6.356e+05 6.038e+06 0.105 48.2 0.04 29.5 ± 7.8 g1124-10 22 7 35 40 2.781e+05 1.390e+06 0.200 11.1 0.02 55.9 ± 23.2 g1124-10 23 5 54 40 1.986e+05 2.145e+06 0.093 17.1 0.14 25.9 ± 12.1 g1124-10 24 15 86 35 6.810e+05 3.905e+06 0.174 31.1 0.00 48.8 ± 13.7 g1124-10 25 10 57 42 3.783e+05 2.157e+06 0.175 17.2 0.00 49.1 ± 16.9 g1124-10 26 8 93 36 3.531e+05 4.105e+06 0.086 32.7 0.02 24.1 ± 8.9 g1124-10 29 7 101 70 1.589e+05 2.293e+06 0.069 18.3 0.00 19.4 ± 7.6 g1124-10 30 21 84 40 8.343e+05 3.337e+06 0.250 26.6 0.01 69.8 ± 17.1 g1124-10 33 0 7 40 0.000e+00 2.781e+05 0.000 2.2 0.02 0.0 ± 74.1 g1124-10 34 17 95 40 6.754e+05 3.774e+06 0.179 30.1 0.04 50.0 ± 13.3 215 1530 4.126e+05 2.936e+06 23.4 for abbreviations, see page 147 area of basic unit = 6.293e-07 cm2 χ2 = 32.061 with 19 degrees of freedom p(χ2) = 3.1% age calculated using a zeta of 392.9 ± 7.4 for cn5 glass ρd = 1.429e+06 cm2 nd = 2241 ρd interpolated between top of can; ρd = 1.388e+06 cm2 nd = 1092 ρd interpolated between bottom of can; ρd = 1.461e+06 cm2 nd = 1149 age dispersion = 27.489% ns / ni = 0.141 ± 0.010 mean ratio = 0.132 ± 0.014 pooled age = 39.3 ± 3.1 ma central age = 38.5 ± 4.0 m sample gc1052-2: apatite 0 -2 +2 69 0 49 2713 1.00 0.75 0.50 0.25 0.00 0.0 0.5 1.0 1.5 2.0 wt% cl 40 30 20 10 0 0 5 10 15 20 track length (µm) pj appendix sample 2 n fr ac tio n ml: 11.71 µm std dev: 2.36 ft age: 38.5 ma 0.0 0.1 0.2 0.25 wt% cl fi ss io ntr ac k ag e (m a) radial plot of single grain ages* distribution of chlorine (cl) contents in apatite grains single grain age vs. weight % cl for individual apatite grains distribution of confined track lengths * see appendix b in geotrack report gc1052 (available online) for details of radial plot construction. colour datapoints indicate wt% cl: dark blue: <0.1%. green: 0.1-0.2%. pale blue: 0.2-0.3%. purple: >0.3%. stratigraphic age 0 50 100 150 200 147 table abbreviations na = number of grid squares counted in each grain ns = number of spontaneous tracks in na grid squares ni = number of induced tracks in na grid squares ratio = ns/ni u (ppm) = uranium content of each grain (= u content of standard glass × ρ i /ρd) cl (wt%) = weight percent chlorine content of each grain ρs = spontaneous track density (ρs) = ns/ (na × area of basic unit) ρi = induced track density (ρi) = ni/(na × area of basic unit) ft age = fission-track age, calculated using equation b.1 area of basic unit = area of one grid square chi squared = χ2 parameter, used to assess variation of single grain ages within the sample p(chi squared) = probability of obtaining observed χ2 value for the relevant number of degrees of freedom, if all grains belong to a single population age dispersion = % variation in single grain ages ns/ni = pooled ratio, total spontaneous tracks divided by total induced tracks for all grains mean ratio = mean of (ns/ni) for individual grains zeta = calibration constant, determined empirically for each observer ρd = track density (ρd) from uranium standard glass (interpolated from values at each end of stack) nd = total number of tracks counted for determining ρd pooled age = fission track age calculated from pooled ratio ns/ni. valid only when p(χ2) >5% central age = alternative to pooled age when p(χ2) <5% geological survey of denmark and greenland. bulletin 10, 33-36 although the structural framework of the subsurface offshore west greenland has been well documented based on comprehensive seismic analysis (cf. dalhoff et al. 2003), the stratigraphy of the region is less well known. the oldest documented sedimentary rocks drilled offshore west greenland are santonian sandstones reached at td in the 6354/4-1 well (fig. 1) although reworked palynomorphs of carboniferous, triassic and jurassic (kimmeridgian) age have been reported from a number of wells in the region. in order to obtain better constraints on the pre-upper cretaceous stratigraphy, a preliminary screening was undertaken to identify inversion structures and erosional canyons where such deeper stratigraphic levels crop out at the seabed (nielsen et al. 2001). sea-floor sampling at selected sites between 62° and 67°n (fig. 1) was undertaken during the summers of 2003 and 2004. other objectives of these cruises were to seek direct evidence of active petroleum systems, to establish further constraints on tectonic and stratigraphic models, and to obtain a better understanding of the neogene and pleistocene history of the region (dalhoff et al. 2005). the most promising seabed features identified by nielsen et al. (2001) were investigated in more detail using a wide range of techniques in order to optimise sampling positions. in 2003, these techniques included echo sounder, side-scan sonar, singlechannel seismic and video inspection before sampling either by dredge, gravity corer, or by video-controlled grab. in 2004, comprehensive data acquisition with a deep-water sparker system was undertaken before sampling by dredge or gravity corer, supplemented by grab samples at selected stations. sampling a total of 19 dredges and five gravity cores, together with three grab samples, were recovered from the eastern and western flank of canyon c1 (fig. 1). on the davis strait high, four dredge and four grab stations were located on the eastern flank and one gravity core was recovered from the western flank (fig. 1). sampling at the seamounts was carried out by dredge on the flanks (six stations) with one grab station at the top of seamount s3 (fig. 1). twenty-six gravity core stations and one piston core station were located in the areas with shallow direct hydrocarbon indications (dhi), inferred from seismic anomalies or from sea-surface slicks recognised on satellite data; the majority were taken within d1 and d2 with the remaining five stations covering d3 and d4 (fig. 1). however, none of the samples produced significant hydrocarbon readings. from a survey on the labrador shelf, mcmillan (1973) reported a positive correlation between sea-floor outcrop and debris composition and concluded that only a small proportion of the recovered rock samples are ice rafted debris (ird). in this study, differentiation between in situ samples and ird continental crust in the davis strait: new evidence from seabed sampling finn dalhoff, lotte m. larsen, jon r. ineson, svend stouge, jørgen a. bojesen-koefoed, susanne lassen,antoon kuijpers, jan a. rasmussen and henrik nøhr-hansen © geus, 2006. geological survey of denmark and greenland bulletin 10, 33–36. available at: www.geus.dk/publications/bull 33 fig. 1. map of labrador sea and davis strait showing sample objectives of the 2003 and 2004 seabed sampling programme referred to in the text. 58˚w 56˚w 54˚w 52˚w 50˚w nuuk maniitsoq c1 s1 d2 c2 s2 d1 c3 s3 d3 d4 s4 hellefisk-1 kangâmiut-1 ikermiut-1 nukik-2 nukik-1 6354/4-1 100 km 66˚n 68˚n 64˚n 62˚n greenland sisimiut hecla high eruption centre maniitsoq high eruption centre davis strait high is based on the number of samples and the lithological homogeneity of the rock types collected at each station. for each dredge and grab station, the samples were classified into three groups: sedimentary rocks (siliciclastic and carbonates), igneous rocks, and precambrian crystalline rocks. quaternary sedimentation understanding of quaternary sedimentation patterns offshore west greenland has greatly improved based on these surveys. gravity cores from the site survey acquired prior to drilling the 6354/4-1 well (fig. 1) and the 37 cores from the seabed sampling programme between 62°n–67°n have formed the basis for a detailed stratigraphic study of the uppermost sediment package (lassen et al. 2005). based on magnetic susceptibility profiles, 14c accelerator-mass-spectrometry (ams14c) dates and core descriptions, it has been possible to relate the succession in the study areas to published event stratigraphic schemes. in general, cores from intermediate water depths (all areas except d2 and the davis strait high, see fig. 1) contain hemipelagic sediments alternating with ird-rich layers. since this pattern is found in all of the studied areas, it is concluded that no resedimentation of the upper sediments has occurred. ordovician and jurassic sediments the identification of dipping reflectors extending up to or near the seabed on conventional seismic lines presented the possibility of sampling sedimentary successions of unknown age at several stations. test sampling (one gravity core) on the eastern flank of canyon c1 (fig. 1) in 2002 revealed reworked middle–upper jurassic and lower cretaceous palynomorphs (h. nøhr-hansen, unpublished data 2004). extensive sampling at this locality was undertaken and a pseudo stratigraphy has been established with the danian–selandian boundary identified on the western flank and the presence of archaean crystalline basement overlain by upper ordovician – lower silurian carbonates and marine mesozoic sediments on the eastern flank. at the davis strait high sampling was undertaken together with a wide range of geophysical investigations indicating the presence of a possible inverted palaeozoic carbonate platform. the assemblage of sedimentary rock types dredged up from the eastern side of the high (fig. 1) is characterised in particular by marine carbonates, which form over 60% of the dredge clasts. other rock types include fragments of sandstone, precambrian crystalline rocks and igneous rocks. the carbonates comprise a range of lithofacies from biomottled lime mudstones and skeletal wackestones (about half of the carbonate clasts from both davis strait high sites) through skeletal, peloidal and intraclastic packstones and grainstones to rare microbial carbonates, including oncolitic and stromatolitic carbonates. partially and wholly dolomitised representatives of these lithofacies are also present. quartz silt and sand are common within the carbonates and fine-grained calcareous or dolomitic sandstones in the sample set are thought to form part of the facies spectrum. the bioclasts within these carbonates include crinoid, brachiopod, mollusc, bryozoan, coral, stromatoporoid, ostracod and trilobite fragments, testifying to the open marine, palaeozoic affinity of the sediments; the facies represented record primarily a low-energy subtidal setting, although high-energy subtidal and lagoonal/intertidal settings are also indicated. organic geochemical analyses of the sedimentary rocks have only revealed two samples with promising results. one sample from canyon c1 and one sample from the davis strait high (fig. 1). the sample from canyon c1 is a laminated upper ordovician mudstone, showing hydrogen index ~400, total organic carbon = 1.3–4% and generative potential (s2-value) that indicates a type 2 source rock with good to excellent potential for oil generation. for a marine carbonate, the biomarker signature shows unusually high proportions of gammacerane, 28-nor-spergulanes (nytoft et al. 2006) and tetracyclic polyprenoids (tpp) and very high proportions of c29 regular steranes. similar characteristics are found in samples of the aleqatsiaq fjord formation in north greenland (christiansen 1989) and in ordovician carbonates from hudson bay and around foxe basin, canada (m.g. fowler, personal communication 2004). the sample from the davis strait high is an ordovician carbonate with oil staining of same age as seen in the source rock from canyon c1. biostratigraphic studies indicate a late ordovician age (maysvillian to richmondian) for davis strait high sample suite, suggesting correlation with lower palaeozoic carbonate successions of the south-east arctic platform and the hudson platform in eastern and north-eastern canada (sanford & grant 2000) and north greenland, and ‘fossilik’ in west greenland (stouge & peel 1979). the correlation with canada and north greenland is supported by the geochemical data. igneous rocks igneous rocks constitute 10–20% of the clast population in most of the dredges, which is a much larger proportion than in the onshore areas. the clasts are typically 5–25 cm in size, rarely > 30 cm across. the majority are massive or vesiculated, fine-grained plagioclase-phyric or aphyric basalts. there are a few oxidised and brecciated samples interpreted as flow tops, and some picrites and gabbros. the clasts vary from angular to rounded and nearly all show abrasion indicative of transportation. notable exceptions are two irregular, 34 very friable rock fragments from seamount s2 (fig. 1) which must be very local. of the total of 195 igneous rock samples, 95 were selected for geochemical analysis. of these, 12 are interpreted to be of precambrian age based on petrography (alteration) and geochemistry, and they are considered to be ird; these include the gabbro samples. the remaining 83 samples can be grouped into relatively few geochemical groups (fig. 2). all the analysed samples from the davis strait high are tholeiitic basalts showing moderate to strong depletion in the most incompatible elements, with ree patterns sloping down to the left (fig. 2a) and low nb/la and nb/lu ratios mostly < 1 (fig. 2b). four samples have increased contents of a number of elements notably si, k, rb, ba, th and pb, and are depleted in nb indicating contamination by continental crustal material (fig. 2b). in the region south of 64°n the majority of samples belong to one geochemical type (the ‘main group’). these rocks are tholeiitic basalts with moderate contents of incompatible elements, ree patterns sloping down to the right (fig. 2a), and nb/la and nb/lu ratios > 1 (fig. 2b). a group of eight samples have relatively flat ree patterns and nb/la and nb/lu ratios close to one. two of these samples show signs of crustal contamination (fig. 2b). a group of six samples are enriched tholeiitic basalts with steeper ree patterns and higher nb contents than the main group. two friable samples of obvious local origin from seamount s2 are alkaline; they are basanites with steep ree patterns and high nb/lu ratios > 10 (fig. 2). other samples include two probably crustally contaminated basaltic andesites and two different rhyolites. radiometric ages have been obtained by the 40ar/39ar step-heating technique for 12 samples. one sample from the flat-patterned group yielded 119 ± 1 ma (early cretaceous), one sample from the davis strait high yielded 63.0 ± 0.7 ma (early paleocene), whereas ten samples from the region south of 64°n yielded ages ranging from 58.0 ± 0.4 ma to 48.1 ± 0.9 ma (late paleocene – early eocene). several lines of evidence support a relatively local derivation for the palaeogene igneous rocks. the high abundance of igneous clasts in the dredges is incompatible with the absence of such rocks in the neighbouring onshore areas. if the basalts in the dredges in the southern region were ice-rafted from the northern onshore volcanic areas in west greenland and baffin island, they should have the same composition and be more frequent in the northern dredges, but this was not observed. the geochemically depleted rocks from the davis strait high are compositionally similar to lava flows outcropping on the sea floor in the same area and sampled in short drill cores (srivastava et al. 1982; m.-c. williamson, unpublished data), making a strong argument for local derivation of these samples. the crustally contaminated samples support the presence of continental crust in the davis strait high, as suggested by keen et al. (1974). the stations south of 64°n are situated immediately south of the large volcanic complexes of the hecla high eruption centre (fig. 1). the lava succession of the hecla high is dated as eocene, and seismic sections show that an eroded and tilted cuesta landscape of lava flows has been exposed subaerially until some time during the pleistocene (sørensen 2005). it is therefore considered most likely that the dredged volcanic rocks from south of 64°n are derived from the hecla high. seamount s2 (fig. 1) is flat-topped with a number of small volcanic cones rising from the platform. the cones probably consist of strongly alkaline rocks as judged from the occurrence of very local samples of basanite. 35 1 10 100 1000 basanite enriched main group flat-patterned depleted 0 1 10 100 nb/la(n) crustally contaminated sa m p le /c h o n d ri te n b /l u (n ) basanite enriched main group flat-patterned depleted rhyolite bas. andesite a b 0.0 0.5 1.0 1.5 la ce pr nd sm eu gd tb dy ho er tm yb lu fig. 2. geochemistry of the dredged igneous rocks. a: rare earth element (ree) patterns for representative samples, chondrite normalised data. the dotted green lines indicate the compositional range of the ‘main group’. b: nb/lu vs nb/la, n means primitive mantle normalised data. normalisation values for both plots from mcdonough & sun (1995). the slopes of the ree patterns indicate that the geochemically depleted basalts from the davis strait high were produced by melting to relatively shallow levels in the spinel stability field in the mantle, i.e. beneath a relatively thin (< 80 km) lithospheric lid. however, the contaminated samples show that they also came into contact with continental crust. the non-depleted and enriched basalts from the hecla high were produced by melting at deeper levels in the garnet stability field in the mantle beneath a thicker lithospheric lid, and the basanites from seamount s2 were produced by the smallest degrees of melting at the deepest levels. the mantle source could in all cases be the proto-icelandic mantle plume. for a more detailed discussion of the igneous rocks see larsen & dalhoff (2006). archaean gneiss two dredges on the eastern side of canyon c1 yielded precambrian gneisses that are considered to be of local origin because of the uniform lithology and unabraded angular morphologies. dating of a gneiss sample by the 206pb/207pb method yielded a well-defined archaean age (2740 ± 150 ma) with no indications of later thermal events. this is compatible with ages derived from the gneisses of the central archaean craton of west greenland, suggesting that the basement gneisses exposed in canyon c1 represent the offshore continuation of this cratonic basement complex. this would extend the known occurrence of archaean basement to nearly 200 km offshore, well into areas that were previously interpreted to comprise oceanic crust (srivastava 1978; roest & srivastava 1989). conclusion and future work the seabed sampling programme has yielded much new information on the early palaeozoic, mesozoic and palaeogene history of the davis strait and labrador sea including the extension of the continental crust into areas earlier described as oceanic crust. continuation of the project with extension of the sampling region to 71°n, west of disko, is planned for the summer of 2006 with collection of additional geophysical data and samples in selected areas. acknowledgement the bureau of minerals and petroleum, nunaoil a/s and encana corporation are thanked for financial support. references christiansen, f.g. 1989: petroleum geology of north greenland. bulletin grønlands geologiske undersøgelse 158, 92 pp. dalhoff, f., chalmers, j.a., gregersen, u., nøhr-hansen, h., rasmussen, j.a. & sheldon, e. 2003: mapping and facies analysis of paleocene – mid-eocene seismic sequences, offshore southern west greenland. marine and petroleum geology 20, 935–986. dalhoff, f., kuijpers, a., nielsen, t., lassen, s., bojesen-koefoed, j.a., larsen, l.m., stouge, s., rasmussen, j.a. & nøhr-hansen, h. 2005: seabed sampling offshore west greenland – new information of inverted palaeozoic and mesozoic basins. gac-mac-cspg-csss joint meeting , halifax, nova scotia, canada, 15–18 may, 2005, abstracts 30, 39 only. keen, c.e., keen, m.j., ross, d.i. & lack, m. 1974: baffin bay: small ocean basin formed by seafloor spreading. american association of petroleum geologists bulletin 58, 1098–1108. larsen, l.m. & dalhoff, f. 2006: composition, age, and geological and geotectonic significance of igneous rocks dredge from the labrador sea and the davis strait. danmarks og grønlands geologiske undersøgelse rapport 2006/43, 73 pp. lassen, s.j., dalhoff, f., kuijpers, a. & nielsen, t. 2005: late quaternary sedimentation patters offshore west greenland based on magnetic susceptibility profiles. danmarks og grønlands geologiske undersøgelse rapport 2005/71, 12 pp. mcdonough, w.f. & sun, s.-s. 1995: the composition of the earth. in: mcdonough, w.f. et al. (eds): chemical evolution of the mantle. chemical geology 120, 223–253. mcmillan, n.j. 1973: surficial geology of labrador and baffin island shelves. in: hood, p.j. (ed.): earth science symposium on offshore eastern canada. geological survey of canada, paper 71-23, 451–469. nielsen, t., andersen, c., jensen, j.b., kuijpers, a., marcussen, c., piasecki, s., sønderholm, m. & rasmussen, r. 2001: geohazard 2001 study offshore west greenland. danmarks og grønlands geologiske undersøgelse rapport 2001/127, 64 pp. nytoft, h.p., lutnæs, b.f. & johansen, j.e. 2006: 28-nor-spergulanes, a novel series of rearranged hopanes. organic geochemistry 37, 772–786. roest, w.r. & srivastava, s.p. 1989: sea-floor spreading in the labrador sea; a new reconstruction. geology 17, 1000–1003. sanford, b.v. & grant, a.c. 2000: geological framework of the ordovician system in the southeast arctic platform, nunavut. in: mccracken, a.d. & bolton, t.e. (eds): geology and paleontology of the southeast arctic platform and southern baffin island, nunavut. geological survey of canada bulletin 557, 13–38. sørensen, a.b. 2006: stratigraphy, structure and petroleum potential of the lady franklin and maniitsoq basins, offshore southern west greenland. petroleum geoscience 12, 221–234. srivastava, s.p. 1978: evolution of the labrador sea and its bearing on the early evolution of the north atlantic. geophysical journal of the royal astronomical society 52, 313–357. srivastava, s.p., maclean, b., macnab, r.f. & jackson, h.r. 1982: davis strait: structure and evolution as obtained from a systematic geophysical survey. in: embry, a.f. & balkwill, h.r. (eds): arctic geology and geophysics. canadian society of petroleum geologists memoir 8, 267–278. stouge, s. & peel, j.s. 1979: ordovician conodonts from the precambrian shield of southern west greenland. rapport grønlands geologiske undersøgelse 91, 105–109. 36 authors’ addresses f.d., l.m.l., j.r.i., s.s., j.b.-k., a.k. & h.n.-h., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: fd@geus.dk s.l., 96 settrington road, london sw6 3ba, uk. j.a.r., geological museum, university of copenhagen, øster voldgade 5–7, dk-1350 copenhagen k, denmark. geological survey of denmark and greenland bulletin 7, 2004, p 33-36 as the drinking water supply in denmark is totally based on groundwater, monitoring of the groundwater quality is extremely important to the danish community. with more than 62% of the total land area under agricultural use the danish government has determined that the entire area is vulnerable to nitrate polution, and therefore the groundwater monitoring programme should cover the entire country. the danish groundwater monitoring programme comprises water supply well monitoring, the groundwater monitoring network and agricultural watershed monitoring (figs 1, 2) and is described on www.groundwater.dk. the programme is part of the national action programme for the water environment and nature, novana (svendsen & norup 2004; svendsen et al. 2005). groundwater quality monitoring is carried out on the basis of data from approximately 6200 public water supply wells. furthermore, a detailed analytical programme is carried out on 1415 well screens from the monitoring network comprising 70 catchment areas, and on 100 shallow screens from the five agricultural watersheds (svendsen et al. 2005). the detailed quality monitoring includes analyses for 97 chemical elements, comprising 26 main elements, 14 heavy metals, 23 organic micro-pollutants and 34 pesticides and metabolites. nitrate the water supply wells generally have long screens and are intended to provide representative information on the distribution of the nitrate content in primary groundwater reservoirs. data from water extraction wells are, however, biased since the wells are intended to ensure production of drinking water with nitrate concentrations below the maximum admissible concentration (mac) of 50 mg nitrate per litre. the groundwater monitoring wells give a more accurate picture of the general nitrate pollution in the danish groundwater. in 1998–2004, mean nitrate concentrations were above the mac limit for drinking water in 16.9% of the wells, whereas about 60% had no nitrate (< 1 mg nitrate per litre). the spatial distribution of nitrate in the groundwater aquifers varies. west of the weichselian glaciation borderline (fig. 1), the outwash plains are dominated by upper, unconfined aquifers overlying deeper, confined quaternary and miocene sands. east of the glaciation borderline the sandy meltwater deposits and the pre-quaternary limestone aquifers are generally covered by clayey till which reduces or prevents nitrate pollution (fig. 3). over the past half century the use of fertilisers in farming has intensified dramatically (fig. 4). groundwater from the monitoring screens has been dated using the cfc (chlorofluorocarbon) content (geus 2004), and demonstrates that the highest nitrate values reflect the increase in the use of fertilisers (fig. 4). continued monitoring will show whether the decrease in the use of fertilisers during the last decade will result in a decrease in the nitrate content, or whether the increasing use of manure will maintain high nitrate levels. preliminary data suggest, however, that since 1979 farmers have adapted their spreading practice for fertilisers and manure such that nitrate pollution is now following a declining trend. 33 groundwater quality monitoring in denmark jens stockmarr geological survey of denmark and greenland bulletin 7, 33–36 (2005) © geus, 2005 < 35 000 35 000 – 350 000 350 000 – 3500 000 water extraction in m3 > 3500 000 50 km10°e 12°e8°e 55°n 57°n fig. 1. extraction of water for domestic consumption from approximately 3000 danish public waterworks in 2001. all waterworks (except one) are based on groundwater extraction. red dashed line indicates the position of the weichselian glaciation borderline. the island of bornholm is shown in the inset map. modified from fraters et al. (2005). nickel, arsenic and other inorganic trace elements the most widespread impact from inorganic trace elements is due to nickel derived from oxidation of pyrite, bravoite and other heavy metal bearing sulphides (fig. 5). massive dissolution of sulphides occurs in areas where large-scale extraction of groundwater has caused lowering of the groundwater table, followed by an influx of oxic groundwater. during dissolution, some of the nickel is adsorbed by contemporaneously precipitated manganese oxide. the trapped nickel is, however, released when the groundwater table rises again, since this leads to reducing conditions and dissolution of the manganese oxides. implementation in 2001 of the eu drinking water directive in danish legislation has led to increased concern with respect to arsenic, for which the mac value was decreased from 50 to 5 µg/l. nine per cent of the monitoring well screens in the groundwater monitoring programme currently exceed 5 µg/l in all samples, while 14% exceed 5 µg/l in at least one sample (fig. 6). the distribution of arsenic is to a large degree controlled by redox conditions, as the solubility of arsenic is about ten times higher in reducing environments compared to oxidising conditions. high arsenic concentrations are mainly found in aquifers underlying clayey sediments. as a result of seven years of groundwater monitoring, the background levels of the 23 inorganic trace elements analysed are well known and are illustrated by the cumulative curves in fig. 7. most curves show a regular distribution, e.g. strontium (sr) and mercury (hg), indicating little or no pollution, whereas the skewed curves for aluminium (al) and nickel (ni) reflect pollution. pesticides and metabolites analytical results from the water supply wells, the groundwater monitoring areas and the agricultural watersheds differ markedly (fig. 8). only agricultural pesticides were detected in water samples collected from young groundwater in the agricultural watersheds, whereas in water samples collected in the groundwater monitoring network other pesticides, such as those used in consolidated areas like urban areas, roads or farmyards have also been found. pesticides and their metabolites are found in more than 25% of all wells (table 1; geus 2004). the most frequently found pesticide group consists of triazines and their metabolites. these compounds are commonly found in both farming and urban areas. in the agricultural watersheds the triazines and their metabolites make up about half of all recorded pesticides and metabolites. in water supply wells the analytical data indicate a very high frequency of 2,6-dichlorobenzamide (bam) findings. twenty-five per cent of the wells contain bam, and 10% have concentrations above the mac value of 0.1 mg/l. bam is a metabolite from dichlobenil and chlorothiamide that were commonly used prior to 1997 as a total herbicide in urban areas, along roads and in farmyards. 34 fig. 2. network of danish national groundwater monitoring areas (catchments) and selected agricultural watersheds covered by the extended monitoring programme. from stockmarr & nyegaard (2004). 12°e10°e8°e 55°n 56°n 57°n 50 km groundwater monitoring areas agricultural watersheds 0 20 40 60 80 100% 0–10 10–20 20–30 30–40 40–50 50–60 60–70 70–80 80–90 ≤ 1 mg/l no3 1–50 mg/l no3 > 50 mg/l no3 d ep th ( m ) n = 894 n = 1 500 n = 1 746 n = 1 413 n = 1 050 n = 711 n = 456 n = 282 n = 149 fig. 3. distribution of nitrate versus depth to well screen in groundwater monitoring and water supply wells. the data cover the period 1998– 2003 and are grouped into three classes according to mean nitrate. from geus (2004). pesticides and metabolites 1998–2003 number % number % number of analysed screens screens with findings findings ≥ 0.1µg/l (mac) monitoring network 1020 437 42.8 153 15.0 agricultural watersheds 75 52 69.3 19 25.3 water supply wells 5515 1443 26.2 355 6.4 table 1. pesticides and metabolites in danish groundwater mac: maximum admissible concentration. from geus (2004). the monitoring network demonstrates high detection rates for pesticides in the upper 40–50 m of wells tested, and a lower number of findings with increasing depth (fig. 8). pollution in shallow groundwater drinking water from 628 dug wells and shallow drilled wells used for single private supply and minor partnership supplies (less than nine families) was investigated in a research project focusing on the youngest groundwater (brüsch et al. 2004). in 35% of the investigated wells, pesticides and their metabolites were recorded with values above the mac value for drinking water (0.1 mg per litre), whereas 11% of wells had more than ten times the mac value. the research project showed that nitrate pollution in 22% of the wells was above 50 mg per litre. the mac level for bacteria was exceeded in 48% of the wells whereas 31% contained coliforme bacteria. in total, 68% of the private supply and minor partnership supply wells delivered undrinkable water. fig. 5. occurrence of nickel in danish water supply wells, 1998–2003. slightly modified from geus (2004). fig. 6. occurrence of arsenic in danish water supply wells, 1998–2003. slightly modified from geus (2004). 1940 1945 1950 1955 1960 1965 1970 1975 1980 1985 1990 1995 2000 0 20 40 60 80 100 120 140 nitrate in groundwater (mg/l) fertiliser (kg n/ha) mean nitrate concentration use of fertilisers per year 0 50 100 150 200 250 300 350 400 oxic zone (o 2 > 1 mg/l) anoxic zone (no 3 > 1 mg/l) fig. 4. nitrate content versus cfc age of groundwater. the annual use of nitrogen fertilisers is shown for comparison. groundwater data from the anoxic zone are corrected on the basis of sulphate content. slightly modified from geus (2004). 12°e10°e8°e 55°n 56°n 57°n 50 km > 20 µg/l ≤ 20 µg/l nickel 12°e10°e8°e 55°n 56°n 57°n 50 km > 5 µg/l ≤ 5 µg/l arsenic 35 36 conclusions during the last two decades many water supply wells have been closed due to pollution; initially this was a result of nitrate pollution, and subsequently also to pesticide pollution. nickel contamination has also closed some wells, and it is feared that arsenic will cause closure of more wells. in some cases the water supply companies have extended their production wells to greater depth instead of closing them, but often this is not a viable solution. due to government policy that drinking-water production should be based on pure groundwater requiring only simple treatment (aeration and iron removal), only few water supply companies have had permission to install advanced water treatment for removal of pesticides or nickel. however, more and more water supplies are expected to need advanced water treatment in the future. nitrate pollution is still the most serious problem because intensive agricultural practices cause leakage of nitrate into groundwater and surface water. while a series of action plans have been introduced to reduce nitrate pollution, much of the young groundwater and most of the surface water in denmark are seriously polluted by nitrate. references brüsch, w., stockmarr, j., kelstrup, n., von platen-hallermund, f. & rosenberg, p. 2004: pesticidforurenet vand i små vandforsyningsanlæg. danmarks og grønlands geologiske undersøgelse rapport 2004/9, 85 pp. fraters, b., kovar, k., willems, w.j., stockmarr, j. & grant, r. (eds) 2005: monitoring effectiveness of the eu nitrates directive action programmes. results of the international monno3 workshop, the hague, the netherlands, 11–12 june, 2003. rivm report 680100002/2005, 290 pp. geus 2004: grundvandsovervågning 1998–2003, 48 pp. københavn: danmarks og grønlands geologiske undersøgelse. stockmarr, j. & nyegaard, p. 2004: nitrate in danish groundwater. in: razowska-jaworek, l. & sadurski, a. (eds): nitrates in groundwater. international association of hydrogeologists, hydrogeology, selected papers 5, 187–199. stockmarr, j., nyegaard, p., larsen, c.l., felding, g. & brüsch, w. 2002: groundwater quality monitoring in denmark. third international conference on water resources and environment research (icwrer), dresden, germany 22–26 july, 2002, proceedings, 165–169. svendsen, l. & norup, b. (eds) 2004: novana. det nationale program for overvågning af vandmiljøet og naturen. programbeskrivelse – del 1. danmarks miljøundersøgelser, faglig rapport fra dmu 495, 45 pp. svendsen, l.m., van der bijl, l., boutrup, s. & norup, b. (eds) 2005: novana. det nationale program for overvågning af vandmiljøet og naturen. programbeskrivelse – del 2. danmarks miljøundersøgelser, faglig rapport fra dmu 508, 128 pp. author’s address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: sto@geus.dk 0.0001 0.001 0.01 0.1 1.0 10 100 1000 10 000 g/l asas crcr cucu ii lili ni hghg cdcd b ba brbr srsrpbpb as cr cu i limo va znhg cd b ba br srpb % 0 20 40 60 80 100 alalal fig. 7. inorganic trace elements found in danish monitoring wells from 1993 to 1999. the skewness of the distribution pattern in the curves for nickel (ni) and aluminium (al) reflects pollution. regular distributions, as e.g. for hg and sr, indicate little or no pollution. slightly modified from stockmarr et al. (2002). fig. 8. occurrence of pesticides and metabolites in water supply and groundwater monitoring wells versus depth to top of well screen, 1998– 2003 as percentage of number of screens investigated. slightly modified from geus (2004). d ep th ( m ) findings above maximum admissible concentrationfindings 0 10 20 30 40 50 60%0 10 20 30 40 50% groundwater monitoring wellswater supply wells 0–10 10–20 20–30 30–40 40–50 50–60 60–70 70–80 80–90 90–100 geological survey of denmark and greenland bulletin 23, 2011, 53–56 53 geological data at the geological survey of denmark and greenland (geus) have been available on the internet for more than 10 years. the first step in making geological data available online was the launch of web access to data from water supply wells (tulstrup 2004). the database is called jupiter, and currently data from more than 260 000 shallow wells are available to the public. figure 1 shows an example of a map from the jupiter database available in a web-browser. the first web access was via a text-based search form which supplied data lists and graphical well reports. in recent years, the interface has been extended with more data, map interfaces and extra functionality. this paper describes this development and illustrates the increasing value of the digital data at geus. in its current form, the jupiter webpage offers: (1) data on wells, geology, water level and groundwater chemistry, (2) free, online danish shallow geological data martin hansen and bjarni pjetursson fig. 1. map interface showing boreholes (dots) and water extraction plants (triangles). the inset map shows the location in south-western denmark. © geus, 2011. geological survey of denmark and greenland bulletin 23, 53–56. open access: www.geus.dk/publications/bull 5454 data about water supply, water abstraction licences, yearly abstraction, exchange of drinking water between waterworks installations as well as drinking-water chemistry and (3) the possibility to download complete data sets in various database formats. the simple text-based search form has been extended with new functions and supplied with different types of map interfaces. data can now be accessed via web map services, web feature services and in google earth format. in 2007, after implementation of a local government reform in denmark, jupiter became the national database for shallow geology, groundwater and drinking water. a data model based on the jupiter database was established and all the data were made available to the public. a set of simple object access protocol web services was launched giving full read-only access to all data in the public domain data model. editing was allowed for the part of the data model that is maintained outside the survey. after the reform of the local government system, tasks involving shallow geology and hydrogeology were divided between (1) the state, which is responsible for hydrogeological mapping and groundwater monitoring, (2) the regions for dealing with soil pollution and remediation and (3) the municipalities (kommuner) for issuing groundwater abstraction licences and checking drinking-water quality. with the demand for geological data at three administrative levels the ability to share knowledge is important. in addition to geological and hydrogeological information from the jupiter database, shallow geophysical data from the survey’s geophysical relational database (gerda) can now fig. 2. map of denmark showing nitrate analyses from groundwater samples. these data are made available through web map services, web feature services and a simple map interface. thirty-five other types of analyses are available in the same form. 55 be used free of charge (møller et al. 2009). databases with hydrological reports and geological models were also established. the reports and models are also available at no cost. at the same time, more than 750 000 documents from the surveys’ old well data archive, and files from well archives held by the former danish counties (amter) were added to the website, all linked to the wells they describe. the data in the jupiter database are in accordance with danish legislation updated by geus and by local authorities. as an example, laboratories must deliver analyses of data set website map download google wms/wfs$ soap# interface earth web services format jupiter yes yes yes yes yes yes gerda yes yes yes no yes no report database yes yes yes no yes no database for geological models yes yes yes no no no groundwater chemistry for 36 selected parameters yes* yes yes* yes yes no $web map services/ web feature services, *part of jupiter, #simple object access protocol. table 1. the databases can be accessed via different interfaces fig. 3. a close-up of the area west of aarhus showing gerda data coverage of different geophysical data types. the inset map shows the location in central denmark. 5656 drinking and groundwater conducted for waterworks directly to the database. when data are entered into the jupiter database, users responsible for drinking-water quality receive notification by e-mail, and after quality control the data are immediately available online. users of the data the jupiter data are used by a wide range of people. by storing data in a central database, the data of one municipality are available not only to the neighbouring municipalities but also to users of drinking water, to the educational system as well as to advisors working for the environmental centres in denmark. different users with different needs and skills call for different user-interface types. at present, data can be accessed through a variety of interfaces. maps and google earth can be used by the public, whereas web map services and web feature services are mainly for professional users (fig. 2). the web services give full reading access through specially designed software and allow editing possibilities for public employees with the appropriate privileges. it is also possible to download complete databases that can be used for complex analyses and to develop geological and hydrogeological models. online shallow geophysical data the geophysical relational database (gerda) contains geophysical data acquired during hydrogeological mapping in denmark over the past decade. the gerda data became available to the public on 1 january 2007. gerda comprises a wide range of geophysical data, geoelectric and geoelectromagnetic data, both raw and processed, reflection seismic data and borehole logging data (fig. 3). inverted 1d and 2d models are included with the geoelectric and electromagnetic data, and the processed sections are available with the seismic data. all information on data acquisition, data processing and inversion procedures can be stored, thus facilitating full reprocessing and inversion of data when required, which makes the inversion and interpretation of data transparent. hydrological reports and geological models geus also hosts a database with hydrogeological reports and a database for geological models. both databases have been developed in cooperation with the danish environmental centres. reports from hydrological surveys are stored in the report database, which was established to allow easy exchange of information between administrators at different levels. reports can be accessed from a search form or from a map interface, if the report is geocoded. the database containing geological models was established to store geological and hydrological models in a toolindependent format. the database is closely connected with the borehole, geophysical and report databases to give the users easy access to reports describing the models and to the data on which the models are based. all models developed by or for the environmental centres during the mapping of the danish groundwater are stored in the model database. concluding remarks most of the danish environmental data are available free of charge and most of them can be found through different danish interfaces (table 1). users can access data through websites, where they can search for specific data sets; they can find data through different types of map interfaces, download parts of databases or complete databases. through web services users can read directly from the database and privileged users can update data belonging to their own administrative unit. the easy access to data makes it easy to share data between different administrative units, between individuals and between consultant companies. private companies, often working for the public administration, have benefitted from the central data storage as they can now access most of the relevant data from one website. in addition, the data are updated and always in the same formats. references møller, i., søndergaard, v.h. & jørgensen, f. 2009: geophysical methods and data administration in danish groundwater mapping. geological survey of denmark and greenland bulletin 17, 41–44. tulstrup, j. 2004: environmental data and the internet: openness and digital data management. geological survey of denmark and greenland bulletin 4, 45–48. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: mh@geus.dk geological survey of denmark and greenland bulletin 20, 2010, 43–46 43 nutrient-poor, low-productive (oligotrophic) soft-water lakes in the atlantic areas of west and north-west europe – the so-called lobelia lakes – are of high conservation value as their low nutrient status favours a particular submerged macrophyte flora with isoetids, which are becoming increasingly rare or threatened due to nutrient enrichment (eutrophication) associated with landuse changes and urbanisation. european union member states have a duty of care, under the habitats directive, to protect the biodiversity of oligotrophic to mesotrophic (moderately productive) standing waters. in denmark the majority of lobelia lakes are located on sandy soils in central and western jylland. these lakes are clear-water ecosystems poor in nutrients and organic carbon and with a unique macrophyte vegetation of predominantly lobelia dortmanna (water lobelia), littorella uniflora (shore-weed) and isoëtes lacustris (quill-wort). severe deterioration of isoetid plant communities is reported from denmark and many other european countries (e.g. arts 2002; pedersen et al. 2006). the isoetids are low and slow growing with relatively poor competitive capabilities. these characteristics make them more sensitive to decreased light levels than other macrophyte groups (middelboe & markager 1997) and consequently also particularly vulnerable to eutrophied and turbid waters. in an ongoing project we are investigating the limnological development of two lobelia lakes in mid-jylland during the last 1000 years (hampen sø and rævsø, situated 2.5 km apart). the hampen sø investigation is part of a geocenter denmark funded project with the title: lake response to climate change during the last 1000 years. in this paper we present the first results from hampen sø with emphasis on changes in the nutrient status of the lake through the last c. 300 years as inferred from diatom and macrofossil analyses. today hampen sø is influenced by nutrient enrichment and has a mixture of two vegetation types that normally belong to two different lake types, namely the lake’s original isoetid vegetation plus species of fast-growing and tall elodeids, the latter being favoured by increased nutrient levels (moeslund 2000). palaeolimnological methods are used to explore the timing and possible causes of the nutrient enrichment. these are of significance for understanding the environmental threat to the lake ecosystem, and for determining its baseline, or reference conditions, defined by the european water framework directive (wfd) as conditions under minimal anthropogenic disturbance. material and methods hampen sø is located in mid-jylland on sandy soils just west of the main stationary line (fig. 1). the lake has a surface area of 76 ha, a mean water depth of 4.3 m, a maximum depth of 13.1 m, and a topographic catchment area of 916 ha (moeslund 2000). the lake is a seepage lake, i.e., a closed lake without natural inlets or outlets. in 2009 a c. 2 m long sediment core with humic, slightly silty and sandy gyttja was retrieved from the lake at a water depth of 9.66 m. the uppermost 1 m of the sediment sequence, which spans approximately the last 1000 years, has been dated by accelerator mass spectrometry (ams) 14c-age determination (one sample) and the 210pb dating method (appleby 2001); here we focus on the period between c. ad 1750 and today. the nutrient status of the lake during this time period is inferred from diatom and macrofossil analyses, and changes in the catchrecent changes in the nutrient status of a soft-water lobelia lake, hampen sø, denmark kaarina weckström, peter rasmussen, bent vad odgaard, thorbjørn joest andersen, tarmo virtanen and jesper olsen 10°e 14°e 55° 57° 100 km sweden denmark germany hampen sø jylland main stationary line fig. 1. map of denmark with the location of the study site, hampen sø. sandy soils dominate west of the main stationary line. © geus, 2010. geological survey of denmark and greenland bulletin 20, 43–46. open access: www.geus.dk/publications/bull 4444 ment landuse are estimated from land classification on two cadastral maps from 1872–74 and 1984–85. diatom data are presented as percentages and 300–400 valves (halves of their siliceous cell walls) were counted per sample. macrofossils are presented as concentrations. chronology the uppermost 28 cm of the lake sediments were dated by 210pb and 137cs assay, estimated from 15 samples spanning the time period 1900–2009. a sample of deciduous leaf fragments found at 117 cm depth was ams 14c-dated with the result cal. ad 910–1020 (range 2 σ). an age–depth curve was constructed for the depth interval 0–117 cm by linear interpolation between the 210pband 137cs-dated sediments and the 14c date at 117 cm. landuse change in the catchment area according to the cadastral maps the most significant landuse change in the hampen sø catchment area since the end of the 19th century is the virtual disappearance of heathland (fig. 2). between c. 1870 and 1980 its area has decreased from c. 24% to less than 1%, while the forested area has increased from c. 36% to 64%. the area of arable land has stayed approximately the same during this time period. the former heathland has been planted with predominantly coniferous trees (pine and spruce) for timber. development of the nutrient status over the last c. 300 years there are two distinct changes in the hampen sø diatom assemblages: in the mid-18th century and around the 1940s (fig. 3). the first change is defined by a clear decrease in planktonic taxa from c. 28% to 8%. variations in the abundance of planktonic species in temperate lakes are often associated with either changing water levels (heinsalu et al. 2008; laird & cumming 2009) or eutrophication (sayer et al. 1999; bennion et al. 2004). cyclotella comensis, an oligotrophic species, dominates the planktonic assemblages prior to the decrease, after which benthic taxa belonging particularly to the genus fragilaria (including staurosira, staurosirella and pseudostaurosira) increase. such a change most likely signifies lowering of the water level. the macrophyte vegetation on the other hand does not show distinct changes, although an increase in carophytes (nitella sp. and chara sp.) can be observed and might suggest that they have been favoured by better light conditions due to shallower water. as the core was collected in the deepest part of the lake, the concentrations of littoral species such as lobelia and isoëtes should be considered with some caution, due to their limited seed and spore dispersal. the first forest plantations in the hampen sø catchment area date back to 1805. at first success was limited until the planting methods were changed in the 1830s. an increase in forest cover will increase evapotranspiration from the catchment area and hence could affect the water level in the lake due to a decrease in the ground water table. we initially assumed that the plantations may have affected the lake levels of hampen sø, however, the water-level decrease indicated by the diatom assemblages clearly occurs before the planting of trees. on the other hand, there is evidence from several studies based on a number of proxies that precipitation decreased in scandinavia during the latter part of the little ice age (e.g. linderholm & chen 2005; de jong et al. 2009). hence the diatom-inferred decrease in hampen sø water levels in the mid-18th century could reflect this suggested change in precipitation. the second clear change in the diatom assemblages occurs around the 1940s, marked by a pronounced increase in the mesotrophic species cyclotella pseudostelligera and fragilaria 1 km 1872–1874 1984–1985 forest field heath wetland lakes landuse fig. 2. landuse in the topographic catchment of hampen sø in 1872–1874 (left) and 1984–1985 (right). 45 crotonensis (up to 23% and 9%, respectively). their increase is followed by the appearance of the eutrophic stephanodiscus parvus in the early 1960s. these taxa indicate increased nutrient concentrations and, with higher phytoplankton productivity, also decreased transparency of the lake water. the charophytes nitella sp. and chara sp., which can occur at water depths of over 10 m, exhibit a decreasing trend from the 1920s (fig. 3), whereas according to recent surveys (moeslund 2000), tall elodeids, which occur in shallower water (such as potamogeton and myriophyllum species), have become more abundant. many elodeids are scarce in sediment records (such as in hampen sø) due to lower seed production and limited dispersal compared to other macrophytes. the reasons for these changes are likely to be the increased nutrient load from the catchment area (intensified field fertilisation and particularly waste waters from a farm and summer houses near the lake, moeslund 2000) which, with increasing phytoplankton productivity, also affect water transparency. concentrations of ephippia (resting eggs) of the water flea genus daphnia began to increase around the same time as the change in macrophyte vegetation is observed with the highest concentrations from the 1960s onwards (fig. 3). daphnia resting eggs are generally produced when conditions deteriorate due to e.g. overcrowding, limited food availability, extreme changes in the environment or increased predation (korhola & rautio 2001). in hampen sø, we interpret the increased numbers of resting eggs as simply an indication of larger daphnia populations due to a general increase in biological production. in addition, more suitable habitats could be represented by the tall elodeids, which function as refuge from fish predation (jeppesen et al. 1997). according to surveys made in hampen sø from 1971 to 1999, nutrient concentrations were high from the beginning of the survey until the mid-1980s (summer-time total phosphorus value c. 70–80 μg l-1 and total nitrogen c. 800–1000 μg l-1), after which concentrations decreased to present-day levels of <30μg l-1 total phosphorous and <600 μg l-1 total nitrogen (moeslund 2000). this decrease in concentrations is attributed to the cessation of waste water effluents (in particular animal manure) from the nearby farm. the decreased nutrient concentrations from the mid-1980s onwards are not reflected in the biota. this could be explained by the sediment nutrient pool which, due to its limited binding capacity (low iron:phosphorus ratio), releases particularly phosphorus back into the water. this is then taken up by phytoplankton and macrophytes (moeslund 2000). st ep ha no dis cu s p ar vu s cy clo te lla p se ud os te llig er a cy clo te lla co m en sis fr ag ila ria cr ot on en sis fr ag ila ria e llip tic a ag g. ac hn an th idi um m inu tis sim um pl an kt on ic sp ec ie s dap hn ia sp ., ep hi pp ia lo be lia d or tm an na , s ee ds n ite lla sp ., o os po re s ch ar a sp ., o os po re s iso et es la cu str is 1760 1800 1840 1880 1920 1960 2000 a ge ( ye ar s a d ) m eg as po re s 5 20 20 5 50020 20 20 40 5 50 150 50 15 30 452510000 0 % concentration (n/100 ml) fig. 3. diatom and macrofossil diagram showing selected abundant species. diatoms (blue) are given as percentages. macrofossils (daphnia ephippia (red) and submerged macrophyte remains (green)) are shown as concentrations (remains per 100 ml fresh sediment). ages based on the constructed age–depth curve are given on the left of the diagram. 4646 it is noteworthy that the abundance of planktonic diatoms slowly begins to increase before the appearance of the mesoand eutrophic diatom taxa (fig. 3). this could suggest slightly increasing water levels from the late 19th century onwards, which may partly have been masked by the marked increase in the planktonic diatom taxa indicating eutrophication. conclusions and future work distinct signs of anthropogenic disturbance in hampen sø can be seen in the 1920s (reflected in macrophyte vegetation and daphnia abundance) indicating the onset of nutrient enrichment. clear indications of eutrophication are evident from the 1960s onwards in all proxies. although water-column nutrient concentrations have decreased since the mid1980s, no change is observed in the biota. this could be attributed to increased internal loading of nutrients from the sediments. compared to the majority of danish lakes, which have been impacted by anthropogenic activities for centuries (bradshaw et al. 2005, 2006), the timing of these changes is surprisingly late. it appears that sandy soils of central and western jylland have been less intensively used for crop cultivation in the past and hence lakes located in such settings are less affected. according to our results, baseline or reference conditions at hampen sø, as defined by the european water framework directive, could be set at the early 1900s. these reference conditions only define the state of the lake before intensified human impact. our data show that climate exerts a notable influence on the groundwater-fed lake and its biota implying that the physical, chemical and biological status of the lake has changed naturally in the past. in the ongoing work the lake’s response to climate change will be explored further. references appleby, p.g. 2001: chronostratigraphic techniques in recent sediments. in: last, w.m. & smol, j.p. (eds): tracking environmental change using lake sediments. 1. basin analysis, coring, and chronological techniques, 171–203. dordrecht: kluwer academic publishers. arts, g.h.p. 2002: deterioration of atlantic soft water macrophyte communities by acidification, eutrophication and alkalinisation. aquatic botany 73, 373–393. bennion, h., fluin, j. & simpson, g.l. 2004: assessing eutrophication and reference conditions for scottish freshwater lochs using subfossil diatoms. journal of applied ecology 41, 124–138. bradshaw, e.g., rasmussen, p., nielsen, h. & anderson, n. j. 2005: mid to late-holocene land-use change and lake development at dallund sø, denmark: trends in lake primary production as reflected by algal and macrophyte remains. the holocene 15, 1130–1142. bradshaw, e.g., nielsen, a.b. & anderson, n.j. 2006: using diatoms to assess the impacts of prehistoric, pre-industrial and modern land-use on danish lakes. regional environmental change 6, 17–24. de jong, r., hammarlund, d. & nesje, a. 2009: late holocene effective precipitation variations in the maritime regions of south-west scandinavia. quaternary science reviews 28, 54–64. heinsalu, a., luup, h., alliksaar, t., nõges, p. & nõges, t. 2008: waterlevel changes in a large shallow lake as reflected by the plankton:periphyton ratio of sedimentary diatoms. hydrobiologia 599, 23–30. jeppesen, e., lauridsen, t.l., kairesalo, t. & perrow, m.r. 1997: impact of submerged macrophytes on fish-zooplankton interactions in lakes. in: jeppesen, e. et al. (eds): the structuring role of submerged macrophytes in lakes. ecological studies 131, 91–114. korhola, a. & rautio, m. 2001: cladocera and other branchiopod crustaceans. in: smol, j.p., birks, h.j.b. & last, w.m. (eds): tracking environmental change using lake sediments. 4. zoological indicators, 5–41. dordrecht: kluwer academic publishers. laird, k.r. & cumming b.f. 2009: diatom-inferred lake level from nearshore cores in a drainage lake from the experimental lakes area, northwestern ontario, canada. journal of paleolimnology 42, 65–80. linderholm, h.w. & chen, d. 2005: central scandinavian winter precipitation variability during the past five centuries reconstructed from pinus sylvestris tree rings. boreas 34, 43–52. middelboe, a.l. & markager, s. 1997: depth limits and minimum light requirements of freshwater macrophytes. freshwater biology 37, 553– 568. moeslund, b. 2000: hampen sø. miljøtilstanden 1971–1999, 103 pp. vejle amt. pedersen, o., andersen, t., ikejima, k., zakir hossain, m.d. & andersen, f.ø. 2006: a multidisciplinary approach to understanding the recent and historical occurrence of the freshwater plant, littorella uniflora. freshwater biology 51, 865–877. sayer, c., roberts, n., sadler, j., david, c. & wade, p.m. 1999: biodiversity changes in a shallow lake ecosystem: a multi-proxy palaeolimnological analysis. journal of biogeography 26, 97–114. authors’ addresses k.w. & p.r., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: kaaw@geus.dk b.v.o., department of earth sciences, university of aarhus, c.f. møllers allé 120, dk-8000 århus c, denmark. t.j.a., department of geography & geology, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. t.v., department of environmental sciences, p.o. box 65, 00014 university of helsinki, finland. j.o., centre for climate, the environment & chronology, archaeology & palaeoecology building, queen’s university belfast, 42 fitzwilliam street, belfast bt9 6ax, uk. geological survey of denmark and greenland bulletin 7, 2004, p 29-32 29 new techniques using computer controlled scanning electron microscopy (ccsem) and laser ablation – inductively coupled plasma – mass spectroscopy (la-icp-ms) have recently been developed at the geological survey of denmark and greenland (geus) to determine source, compositional variation and sedimentary pathways of sandstones. these new timeand cost-efficient methods are highly applicable in petroleum and mineral exploration. this paper illustrates how the provenance and variability of miocene titanium-rich sands in western and central jylland have been investigated, but the methods are presently also used offshore the faroe islands and in east and west greenland. ccsem and la-icp-ms utilise simple sample preparation methods, are relatively rapid and less expensive than conventional methods and yield more information. heavy mineral (titanium) exploration techniques exploration for economic concentrations of valuable heavy minerals, particularly ti-bearing minerals, necessitates knowledge of their compositional variation and the sand transport processes into the sedimentary basin where they are found (morton & hallsworth 1999; morton et al. 2004). conventional provenance studies employ methods such as sensitive high resolution ion microprobe (shrimp) and electron microprobe analysis (empa) that require time-consuming and often tedious sample preparation, are thus expensive, and are therefore not suitable as a standard exploration technique. as part of a provenance study of ilmenite-rich sands occurring in the miocene of western denmark carried out at geus, it was the aim to develop a more cost-efficient standard analytical tool that could integrate bulk rock chemostratigraphic data, mineral compositional data and age information. this study describes the use of ccsem and la-icp-ms as alternative techniques, capable of providing detailed information relevant to sand provenance more rapidly and less expensively than conventional methods. general geological setting the coastline of the proto-north sea had a nw–se trend across central jylland during the miocene (fig. 1). the coastal regions were sourced by large rivers draining the western part of the scandinavian shield during the early miocene. a distinct change to a more westerly transport direction occurnew methods in provenance studies based on heavy minerals: an example from miocene sands in jylland, denmark christian knudsen, dirk frei,thomas rasmussen, erik s. rasmussen and roger mclimans geological survey of denmark and greenland bulletin 7, 29–32 (2005) © geus, 2005 250 km 350–420 ma 900–1150 ma 1200–1500 ma 1500–1770 ma 1650–1855 ma 1800–1900 ma archaean denmark baltic sea north sea sedimentary cover rocks jylland deltaic and continental deposits beach deposits lagoonal deposits marine deposits stauning 50 km 1 2 3 4 1 sønder omme2 isenvad4 store vorslunde3 8°e 16°e 24°e 32°e 40°e 58°n 66°n fig. 1. left: palaeogeographic map showing distribution of environments in jylland during deposition of the odderup formation (middle miocene; from rasmussen 2004). above: map showing basement ages found in scandinavia and the baltic area (modified from gaál & gorbatschev 1987). red during the mid-miocene (rasmussen 2004). the climate was warm temperate to subtropic. heavy mineral enrichment has been found in beach sands of what were barrier islands (odderup formation and bastrup sand) and in offshore storm sand deposits (stauning sand, figs 1, 2; knudsen 1998; fisher 2003). major and trace element geochemistry during exploration for heavy mineral accumulations, a large number of drill holes were completed in western and central jylland. ilmenite-bearing sand samples (bulk rock) from the drill holes were analysed by x-ray fluorescence (xrf). in the vorslunde area, the analytical results show that the finegrained micaceous sand belonging to the stauning sand (rasmussen 2004) has a distinct geochemical signal compared to the overlying medium-grained quartz sand of the odderup formation (fig. 3). the fe/ti ratio (fig. 3a) is higher in the stauning sand compared to the odderup formation as iron is leached out of the ilmenite during alteration of the sediment (weibel 2003). the higher fe/ti ratio in the stauning sand indicates that it has suffered less leaching than the odderup formation. the v/ti ratio (fig. 3b) is higher in the stauning sand than the odderup formation, indicating a higher vanadium content in the ilmenite of the stauning sand. as vanadium is less likely to be leached out of the ilmenite than iron, this may reflect a difference in the composition of ilmenite in the source areas for the sands in the two units. the y/ti ratio (fig. 3c) is higher in the stauning sand than the odderup formation indicating that the stauning sand has a different heavy mineral assemblage; presumably a higher content of primary monazite and xenotime which are the main carriers of yttrium. this may indicate a different ratio between monazite and xenotime in the source regions of the two units. computer controlled scanning electron microscopy the ccsem method, developed and applied in provenance analyses at geus (frei et al. 2005), provides the following: 1. modal abundances of the heavy minerals. 2. major element composition of the individual grains. 3. grain-size and grain-shape data for the grains. 4. maturity of the ti-minerals. for ccsem analysis, the bulk sample and 500 to 1500 grains of the heavy mineral fraction obtained by heavy liquid separation are embedded in epoxy such that no grains touch each other. 30 odderup fm arnum fm gram fm bastrup sand stauning isenvad vorslunde stauning sand sdr. omme sw ne lithostratigraphy marine silt and clay fluvial and marine sand brackish water silt and clay coal hodde fm fig. 2. simplified miocene lithostratigraphy of the study area with position of sample localities indicated (modified from rasmussen et al. 2004). 0 1 2 3 4 5 a b c 0 2 4 6 8 10 % tio 2 0 2 4 6 8 10 0 20 40 60 80 0 20 40 60 0 2 4 6 8 10 stauning sand odderup formation stauning sand odderup formation stauning sand odderup formation % tio 2 % tio 2 % f e 2 o 3 p p m v p p m y fig. 3. xrf analyses of sand from the vorslunde area. for location, see fig. 1. modal abundances of the heavy mineral suite the chemical composition of the grains is detected as well as such physical parameters as aspect ratio, circularity, perimeter, length and number of grains counted. this information is used to calculate the modal abundances and grain-size distributions of the detected minerals. in fig. 4, the modal abundance of heavy minerals in the odderup formation is compared to the underlying stauning sand. the odderup formation is dominated by stable heavy minerals such as ilmenite, rutile and zircon, whereas the stauning sand contains a large proportion of mafic silicates vulnerable to dissolution during alteration (weibel 2003), indicating a higher maturity of the odderup formation sands. titanium content of the titanium minerals the tio2-distribution in the ti-mineral fraction (i.e. ilmenite, rutile and leucoxene) of a sediment is a very sensitive indicator of its maturity. the continuous alteration of primary ilmenite (with an average stoichiometric tio2-content around 50 wt%) to leucoxene by leaching of iron, leads to a significant increase in the tio2-grade of the ti-mineral fraction in a mature sediment (e.g. weibel 2003). since the ccsem analysis yields the average chemical compositions of all minerals present in a given sample, the tio2-distribution of the ti-mineral fraction can easily be calculated without the need for time-consuming mineral separation. the average tio2-content of around 68 wt% observed in the odderup formation (fig. 5) is indicative of a mature sediment. in contrast, the majority of the data for the stauning sand scatters around the normal tio2-content of 50 wt% (fig. 5), characteristic of primary, unleached ilmenite, indicating a less mature sediment. this supports the interpretation of fig. 3a mentioned above. garnet composition the compositional variability of the detrital garnet fraction in sediments is also a very useful provenance indicator (e.g. morton 1985; morton & hallsworth 1994). garnet compositions are normally determined by point-counting a representative number of grains using electron microprobe analysis. using ccsem, information on the chemical variability of the garnet fraction in a sample is acquired concomitantly with the titanium mineral data during routine operation, i.e. no further sample preparation or analytical steps are required. laser ablation – inductively coupled plasma – mass spectrometry for la-icp-ms age determinations, zircons are separated from the bulk samples using conventional heavy liquid and magnetic separation methods. the final separation step is 31 odderup fm (sønder omme) odderup fm (vorslunde) stauning sand (stauning) stauning sand (vorslunde) ti-magnetite ilmenite leucoxene rutile zircon garnet kyanite-sillimanite -staurolite mafic silicates fig. 4. modal abundances in the heavy mineral fraction of odderup formation sand and stauning sand determined by ccsem. for location, see fig. 1. 30 38 46 54 62 70 78 86 94 0 10 20 30 m in er al w t% 0 10 20 30 m in er al w t% odderup formation stauning sand tio 2 (wt%) 30 38 46 54 62 70 78 86 94 tio 2 (wt%) fig. 5. distribution of tio2-content of the ti-mineral fraction in the odderup formation and the stauning sand at vorslunde, determined by ccsem. the higher tio2-content in the quartz sand reflects its higher maturity. for location, see fig. 1. made by hand-picking individual zircon grains using an optical microscope. for analysis, the individual detrital zircon grains are mounted on adhesive tape. the la-icp-ms dating is based on measuring the 207pb-206pb ratios in the zircon, and the data are in good agreement with data from shrimp analyses (frei et al. 2005). the different provenance of the various miocene sandy units is also reflected by their detrital zircon age distribution data (fig. 6) . the bastrup sand and the odderup formation are characterised by a bimodal distribution of the ages, with peaks around 1100 and 1600 ma, whereas the stauning sand has a broad distribution with a single peak around 1500 ma and a scattering of ages between 2500 ma and 3000 ma. when compared to the map showing the distribution of ages in the scandinavian shield and the baltic area (fig. 1), a north and north-easterly provenance for the bastrup sand and the odderup formation is likely, whereas the stauning sand probably has a more easterly source, where older rocks are more common. however, the bastrup sand and the odderup formation are coarser grained, and were not deposited in the same environment. there is a possibility that the differences are due to sveco-norwegian zircons being larger and thus being less frequent in the fine-grained sediments of the stauning sand. acknowledgement the results presented here are a part of a study carried out with financial support from dupont titanium technologies. references fisher, t. 2003: georadar investigations and sedimentology of miocene heavy mineral deposits, central jylland, 178 pp. unpublished m.sc. thesis, university of copenhagen, denmark. frei, d., rasmussen, t., knudsen, c., larsen, m., whitham, a. & morton, a. 2005: linking the faroese area and greenland: new methods and techniques used in an innovative, integrated provenance study. annales societatis scientiarum færoensis supplementum 43, 1–7. gaál, g. & gorbatschev, r. 1987: an outline of the precambrian evolution of the baltic shield. precambrian research 35, 15–52. knudsen, c. 1998: heavy mineral exploration in miocene sediments, jylland. danmarks og grønlands geologiske undersøgelse rapport 1998/45, 44 pp. morton, a.c. 1985: a new approach to provenance studies: electron microprobe analysis of detrital garnets for middle jurassic sandstones of the northern north sea. sedimentology 32, 553–566. morton, a.c. & hallsworth, c.r. 1994: identifying provenance-specific features of detrital heavy mineral assemblages in sandstones. sedimentary geology 90, 241–256. morton, a.c. & hallsworth, c.r. 1999: processes controlling the composition of heavy mineral assemblages in sandstones. sedimentary geology 124, 3–29. morton, a.c., hallsworth, c. & chalton, b. 2004: garnet compositions in scottish and norwegian basement terrains: a framework for interpretation of north sea sandstone provenance. marine and petroleum geology 21, 393–410. rasmussen, e.s. 2004: stratigraphy and depositional evolution of the uppermost oligocene–miocene succession in western denmark. bulletin of the geological society of denmark 51, 89–109. rasmussen, e.s., dybkjær, k. & piasecki, s. 2004: the billund delta: a possible new giant aquifer in central jutland. geological survey of denmark and greenland bulletin 4, 21–24. weibel, r. 2003: alteration of detrital fe-ti oxides in miocene fluvial deposits, central jutland, denmark. bulletin of the geological society of denmark 50, 171–183. authors’ addresses c.k., d.f., t.r. & e.s.r., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ckn@geus.dk r.mcl., dupont titanium technologies, experimental station, e352/217, route 141 and henry clay, wilmington, de 19808, usa. 32 n u m b er s o f zi rc o n s n u m b er s o f zi rc o n s 0 5 10 15 20 0 0 5 10 15 20 1000 2000 3000 million years 0 1000 2000 3000 million years 0 1000 2000 3000 million years 0 1000 2000 3000 million years odderup formation, vorslunde, quartz sand 123 zircon grains stauning sand, sønder omme, mica sand 181 zircon grains bastrup sand, isenvad, quartz sand 118 zircon grains stauning sand, stauning, mica sand 128 zircon grains 0 5 10 15 20 0 5 10 15 20 fig. 6. age distribution among detrital zircons from miocene sands in jylland. note the bimodal distribution of zircon ages in the odderup formation and the bastrup sand in contrast to the unimodal and generally older ages in the stauning sand (see text for discussion). geological survey of denmark and greenland bulletin 6, 31-49 31geological survey of denmark and greenland bulletin 5, 31–49 © geus, 2004 maximum middle jurassic transgression in east greenland: evidence from new ammonite finds, bjørnedal, traill ø peter alsen and finn surlyk with an appendix by john h. callomon: description of a new species of ammonite, kepplerites tenuifasciculatus n. sp., from the middle jurassic, lower callovian of east greenland a middle – lower upper jurassic sandstone-dominated succession, more than 550 m thick, with mudstone intercalations in the middle part is exposed in bjørnedal on traill ø, north-east greenland. a number of ammonite assemblages have been found, mainly in the mudstones. they indicate the presence of the lower callovian cadoceras apertum and c. nordenskjoeldi chronozones. the mudstones represent northern wedges of the fossilbjerget formation hitherto known only from jameson land to the south. in bjørnedal they interfinger with sandstones of the pelion and olympen formations. the presence of the fossilbjerget formation in this region indicates complete drowning of the middle jurassic sandstone-dominated pelion formation during maximum middle jurassic transgression. a new species, kepplerites tenuifasciculatus, is described in the appendix by j.h. callomon. the holotype and paratype are from jameson land, east greenland, but the species is also found in bjørnedal, traill ø, north-east greenland. keywords: ammonites, fossilbjerget formation, middle jurassic, kepplerites tenuifasciculatus callomon, pelion formation geological institute, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: petera@geol.ku.dk the mesozoic deposits of the traill ø area in northeast greenland were studied by donovan (1953) during lauge koch’s east greenland expeditions (fig. 1). middle jurassic exposures are few and scattered and fossils are scarce, allowing only a few tie-points to the biostratigraphically well-dated succession in jameson land to the south (callomon 1993). a succession of dark micaceous, silty mudstones in an otherwise sandstone-dominated succession in the bjørnedal valley, south-east traill ø was described by donovan (1953; fig. 2). only a few poorly preserved and unidentifiable ammonite fragments were recovered. hence the exact stratigraphic position and age of the mudstones were unknown and have remained as such until reexamination of the locality and intensive search for ammonites was carried out during fieldwork in the traill ø area in 1996 (alsen 1998). new finds of ammonites indicate an early callovian age (cadoceras apertum chronozone) of the mudstones (fig. 3), which are referred to the fossilbjerget formation, previously known only from jameson land (surlyk et al. 1973). this paper presents the new finds of ammonites in the bjørnedal area and the implications of these for the sequence stratigraphic interpretation of the basin. the traill ø area forms the northwards continuation of the jurassic jameson land basin within the east greenland rift basin. huge amounts of sand-dominated sediments were introduced into the basin mainly from the north in middle jurassic times and were transported southwards by marine currents along the basin axis. a geus bulletin no 5.pmd 29-10-2004, 11:1431 32 thick succession of middle jurassic marine sediments, mainly sandstones, was deposited in northern jameson land and progressively thins distally towards the south. the proximal areas in traill ø acted as a bypass area for much of late middle jurassic time and the succession is generally thinner compared with the more distal areas in jameson land (surlyk 2003). the south-easternmost part of traill ø constituted an eastwards-tilted fault block, in contrast to other jurassic fault blocks in east greenland which are tilted westwards (donovan 1953; carr 1998; surlyk 2003). a detailed analysis of the middle jurassic succession in the vælddal area close to the crest of the eastwardstilted block was recently undertaken by carr (1998; fig. 1). he demonstrated that as a result of synsedimentary movements of the vælddal fault during middle jurassic time the sedimentary evolution of the bjørnedal area differed somewhat from that of other parts of the region. the middle jurassic succession in this block shows palaeocurrent directions to the east in contrast to the axial, southwards-oriented palaeocurrent directions seen elsewhere. the sandstone/mudstone ratio also decreases and the succession shows proximal to distal facies changes in the same direction (vosgerau et al. 2004, this volume). stratigraphy the lithostratigraphical scheme for the jurassic of east greenland was erected by surlyk et al. (1973), surlyk (1977, 1978) and birkelund et al. (1984). the scheme is under revision and many units have been raised in rank (see surlyk 2003, fig. 5). in the traill ø area, the middle – lower upper jurassic succession now includes the bristol elv, pelion, fossilbjerget and olympen formations in ascending order (fig. 4). the bristol elv formation is probably not exposed in bjørnedal. it is of pre-late bajocian but probably still middle jurassic age and includes fluvial pebbly sandstones and thin coaly shales (therkelsen & surlyk 2004, this volume). the pelion formation consists of shallow-marine sandg r ee n la n d ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ geographical society ø traill ø jameson land 24°w 23°w 22°w 72°45′n 72°30′n 72°15′n n 10 km kong oscar fjord mountnorris fjord fig. 2 vælddal fault vælddal bjørnedal g re en la nd palaeogene intrusives palaeogene extrusives cretaceous jurassic triassic upper permian carboniferous devonian pre-devonian fault 73°00′n fig. 1. simplified geological map of the traill ø area. location of study area (fig. 2) is indicated. geus bulletin no 5.pmd 29-10-2004, 11:1432 33 stones with bivalves, belemnites, ammonites and plant fragments. the formation rests on triassic strata or on the bristol elv formation. it is overlain by, or interfingers at the top with, offshore, dark, micaceous, silty mudstones of the fossilbjerget formation. a regressive wedge in the top part of the pelion formation interrupts the overall backstepping pelion–fossilbjerget couplet, and belongs to the parnas member of the pelion formation (surlyk 2003). the fossilbjerget formation is overlain by the lower upper jurassic olympen formation, which also consists of shallow marine sandstones themselves quite similar to the sandstones of the pelion formation below, making recognition of the olympen formation difficult. the middle jurassic succession in the bjørnedal area is poorly exposed. the area described here corresponds to the ‘northern part of bjørnedal’ of donovan (1953). an intensive search for ammonites was made in dark, silty, micaceous mudstones exposed in small sections along the bjørnedal river, locality 1, and in a small section in a gully at locality 2 (figs 2, 5). three ammonite assemblages were recovered; two ammonite species have been identified, one of them previously known from jameson land but hitherto undescribed. the description of the middle jurassic succession in bjørnedal is based mainly on donovan (1953), since fieldwork was focused solely on the mudstones in its middle part. the succession is subdivided into five lithological units, a–e in ascending order (figs 4, 5). unit a belongs to the pelion formation and consists of about 200 m of sandstones containing cranocephalites, which suggests correlation with the upper bajocian c. pompeckji chronozone (fig. 3) and indetermilycett bjerg steenstrup bjerg 834 1013 bjørnedal brede gletscher barrikade gletscher mountnorris fjord land ice-covered sea river 2 km n 2 3 1 lo w er c al lo vi an ba th on ia n ba jo ci an orbis zigzag parkinsoni garantiana subfurcatum zones calloviense nw european/ subboreal province boreal province koenigi herveyi discus *nordenskjoeldi (parnas mb, wedge of upper pelion fm) *apertum (fossilbjerget fm) calyx variabile cranocephaloide ishmae greenlandicus arcticus *pompeckji (pelion formation) * zones identified in the bjørnedal area indistinctus borealis tenuiplicatus progracilis subcontractus morrisi hodsoni m id dl e ju ra ss ic humphriesianum fig. 2. map of the bjørnedal area, traill ø showing locations of sections and ammonite localities. 1, locality 1 with section jth15/96; 2, locality 2 with section jth-14/96; 3, locality 3. altitude of mountains in metres. fig. 3. scheme of standard ammonite zones in the upper bajocian – lower callovian, middle jurassic. due to faunal provincialism close correlation between zones in east greenland (right column) and the nw european/subboreal province (left column) is not possible. the east greenland ammonite zonation represents a secondary standard chronostratigraphical scheme for the boreal province. modified from callomon (2003). geus bulletin no 5.pmd 29-10-2004, 11:1433 34 parnas member, representing the last regressive tongue of the overall backstepping pelion formation. unit d consists of about 40–50 m of black, silty mudstones with intercalated sandstones and indeterminable ammonites. it belongs to the upper part of the fossilbjerget formation. a 50 m core (ggu 303124) drilled at locality 2 (fig. 2) through most of this unit records a development from offshore black mudstones with occasional storm sandstone beds, 5–10 cm thick, into offshore transition zone mudstones with an upwards increasing content of fine-grained sandstone (fig. 6). at the top, a syenite sill, c. 3 m thick, occurs between the mudstones and the base of the sandstones of the overlying lower upper jurassic olympen formation (unit e). this formation consists of about 220 m of apparently massive sandstones interbedded with irregular beds of dark micaceous mudstones with traces of plants (donovan 1953). it is overlain by upper jurassic black mudstones of the bernbjerg formation containing late oxfordian ammonites (amoeboceras serratum sowerby 1813; price & whitham 1997). discussion the sequence stratigraphical development of the middle jurassic succession of jameson land was interpreted by surlyk (1990, 1991), surlyk et al. (1993) and in more detail by engkilde & surlyk (2003). it is difficult to trace or correlate key surfaces from jameson land across kong oscar fjord to the traill ø area, in part because of the rather limited and generally poor outcrops on traill ø. in jameson land, the base of the pelion formation appears to be almost isochronous everywhere and of early late bajocian age (cadoceras borealis chron; fig. 3). the formation is overlain by the fossilbjerget formation, with a strongly diachronous boundary younging sourcewards towards the north. this reflects the overall backstepping of the sandy pelion system caused by a combination of increasing rifting and eustatic sea-level rise. during maximum transgression in the early–middle callovian, deposition of offshore fossilbjerget mudstones reached its northernmost extent. the presence of the fossilbjerget formation in bjørnedal, traill ø, extends the known distributional area of the formation northwards by about 100 km. this reflects drowning of the marine sanddominated depositional systems of the pelion formation during the middle jurassic at maximum transgression. a sandstone wedge of the parnas member (pelion formation) is intercalated in the fossilbjerget forma100 m a b c d e a. serratum chronozone (upper oxfordian) c. pompeckji chronozone (upper bajocian) pe lio n fo rm at io n o ly m pe n fo rm at io n fo ss ilb je rg et f m fo ss ilb je rg et f m be rn bj er g fm pa rn as m b (p el io n fm ) black mudstones c. apertum chronozone (lower callovian) c. nordenskjoeldi chronozone (lower callovian) unexposed, mainly sandstones silty shales sandstones fig. 4. schematic section of the middle jurassic succession in bjørnedal, based on donovan (1953) and new data. nate plant remains (donovan 1953). the unit is exposed north-west of the mouth of the valley on the north-eastern slope of mount lycett bjerg facing mountnorris fjord (fig. 2). the strata dip towards the south-east and disappear beneath the succession exposed in bjørnedal. the base of the unit is not exposed. unit b belongs to the fossilbjerget formation and consists of 25–30 m of soft, black, micaceous, silty mudstones with pyritic concretions and indeterminate plant impressions. the species of kepplerites found indicates an early callovian age (cardioceras apertum chron; fig. 3). unit c is a mainly scree-covered interval, about 60 m thick, probably consisting mostly of sandstones. finds of cadoceras sp. in loose blocks in the scree indicate an early callovian age (cardoceras nordenskjoeldi chron; fig. 3). the unit belongs to the geus bulletin no 5.pmd 29-10-2004, 11:1434 35 lycett bjerg sw ne bernbjerg formation olympen formation fossilbjerget formation parnas member fossilbjerget formation 2 3 1 tion and represents the last, early callovian, regressive tongue of the overall backstepping pelion formation. the fossilbjerget formation is overlain by the prograding sandstones of the olympen formation. systematic palaeontology specimens with mguh numbers are stored at the geological museum, university of copenhagen, denmark. superfamily stephanocerataceae neumayr 1875 family kosmoceratidae haug 1887 genus kepplerites neumayr & uhlig 1892 kepplerites cf. tenuifasciculatus callomon 2004 plate 1, fig. 1 material. one fragmented specimen (mguh 25762 from ggu 429773). horizon. the specimen was found loose in unit b at locality 1 (figs 2, 4, 5). description. only two-fifths of the last whorl is preserved, as an imprint in silty mudstone. the fragment, a macroconch, comprises most of the body chamber of an adult specimen with an estimated maximum diameter of 130 mm. its umbilical seam is uncoiling and the ribs near the aperture are strongly projected forward. primary ribs are coarse and strongly curved. each primary rib gives rise to four secondary ribs, which persist to the venter. secondary ribbing is very fine and dense. secondary ribs are almost straight but seem to curve slightly backwards near the venter giving an overall sinuous appearance of the primary and secondary ribs. comparison. the species is consistently more densely, finely ribbed than any other species of kepplerites from east greenland, with little, if any, modification of the primary ribbing in the adult stage. age and distribution. k. tenuifasciculatus occurs in central jameson land and on traill ø. in central jameson land it is found at a level above k. traillensis (faunas 24–26 of callomon 1993). k. traillensis is thought to correlate closely with the horizon of k. keppleri in europe, which defines the base of the callovian. k. tenuifasciculatus is thus of early callovian age (c. apertum chron; fig. 3). fig. 5. view of bjørnedal towards the north-west. the succession exposed on the southern flank of lycett bjerg is subdivided into lithostratigraphical units. 1–3, indicate localities (fig. 2). the fossilbjerget formation section in the centre of the photograph is 40– 50 m thick. (photo: s. piasecki). geus bulletin no 5.pmd 29-10-2004, 11:1435 36 family cardioceratidae von siemiradzky 1891 genus cadoceras fischer 1882 cadoceras sp. plate 1, figs 2a–c, 3 material. several fragments preserved as imprints in black shale (mguh 25763–766 from ggu 429771; ggu 429772). the assemblage includes well-preserved small specimens of microconchs, some of which are almost complete (plate 1, figs 2a–c). horizon. the assemblage was collected at a small exposure in the streambed of the bjørnedal river in bjørnedal, traill ø. the succession was not measured, but belongs to unit b (fig. 4). description. the fragments seem to be those of macroconchs. they are densely ribbed with relatively coarse, strong primaries and secondaries. size cannot be estimated. the microconchs are small and evolute with dense, very fine ribbing. ribbing becomes somewhat coarser near the final peristome. the maximum diameter is on average about 16 mm. age. the specimens are early cadoceras (j.h. callomon, personal communication 1997) probably from the c. apertum or c. nordenskjoeldi chronozones (fig. 3; faunas 27–30 of callomon 1993). the assemblage thus indicates an early callovian age. 0 10 20 30 m te te ph th ? pa te te te ch ch te rh ? te ch te ch o ly m pe n fo rm at io n fo ss ilb je rg et f or m at io n degrees of bioturbation te terebellina isp. ph phycosiphon isp. th pa palaeophycus isp. ch chondrites isp. rh rhizocorallium isp. thalassinoides isp. parallel lamination sandstone siltstone wavy lamination cross-lamination 50 40 sand cl si f m c gr fig. 6. sedimentological log of core ggu 303124, measured by j. therkelsen and p. alsen. for location, see figs 2, 5. geus bulletin no 5.pmd 29-10-2004, 11:1436 37 cadoceras cf. nordenskjoeldi callomon & birkelund 1985 plate 1, fig. 4a–c 1904 olcostephanus neumayr (?simbirskites pavlow & lamplugh) nov. sp. – madsen, p. 195, plate 10, fig. 2. cf. 1985 cadoceras nordenskjoeldi n. sp. callomon & birkelund, p. 84, pl. 1, fig. 4; pl. 4, figs 1–6. material. several fragments (mguh 25767–769 from ggu 429768). the fragments are small and crushed macroconchs and are poorly preserved in hard, reddish sandstone. complete specimens and microconchs have not been identified. horizon. the assemblage was found loose in unit c at locality 3 between the mudstones exposed at localities 1 and 2 (figs 2, 4, 5). description. the inner whorls are finely and densely ribbed and seem to be evolute. the outer whorls are evolute, with very coarse and blunt primary and secondary ribbing. the whole assemblage seems to consist of fairly slim variants. comparisons. the specimens are referred to as c. cf. nordenskjoeldi, which has the above-mentioned characteristic style of ribbing. this is supported by their stratigraphical position above k. tenuifasciculatus. age and distribution. this is the first find of the species outside thefossilbjerget and olympen areas in central jameson land. c. cf. nordenskjoeldi indicates an early callovian age (c. nordenskjoeldi chron; fig. 3). cadoceras sp. indet. plate 1, fig. 5 material. two specimens and fragments (mguh 25770 from ggu 429767; ggu 429769–770). one specimen is a poorly preserved macroconch (plate 1, fig. 5); the other is a poorly preserved microconch. both are imprints in mudstones. horizon. the ammonites were collected at locality 2 from level 29–30.6 m in section jt-14 (fig. 2) in mudstones of unit d (figs 2, 4, 5). description. the macroconch is medium-sized and evolute, and comprises most of a whorl of the bodychamber. ribbing is dense and relatively strong up to the aperture. primaries divide into secondaries on the middle of the flank. the specimen shows no adult modification and could be a juvenile. maximum diameter measures 61.5 mm. the other specimen, probably a microconch, is small and relatively evolute. ribbing is dense and sharp with primaries and secondaries. maximum diameter is 30 mm. age. the assemblage is too poorly preserved for determination to specific level. the specimens are possibly early cadoceras of early callovian age (j.h. callomon, personal communication 1997; fig. 3). acknowledgements we acknowledge generous support by the danish research councils (project: ‘resources of the sedimentary basins of north and east greenland’) and thank jette halskov for drafting, jens therkelsen, who provided the measured sections, john h. callomon and lars stemmerik for reading the manuscript critically, and walter k. christiansen and g. bloos for useful reviews. references including references cited in appendix alsen, p. 1998: middle jurassic ammonite biostratigraphy in the traill ø region, central east greenland. abstracts – 23rd nordiske geologiske vintermøde århus, denmark, 1998, 17 only. birkelund, t., callomon, j.h. & fürsich, f.t. 1984: the stratigraphy of the upper jurassic and lower cretaceous sediments of milne land, central east greenland. bulletin grønlands geologiske undersøgelse 147, 56 pp. buckman, s.s. 1922: type ammonites 4, 67 pp., pls 267–422. london: wheldon & wesley. callomon, j.h. 1993: the ammonite succession in the middle jurassic of east greenland. bulletin of the geological society of denmark 40, 83–113. callomon, j.h. 2003: the middle jurassic of western and northern europe: its subdivisions, geochronology and correlations. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 61–73. callomon, j.h. 2004: description of a new species of ammonite, kepplerites tenuifasciculatus n. sp., from the middle jurassic, lower callovian of east greenland. appendix in: geus bulletin no 5.pmd 29-10-2004, 11:1437 38 alsen, p. & surlyk, f.: maximum middle jurassic transgression in east greenland: evidence from new ammonite finds, bjørnedal, traill ø. in: stemmerik, l. & stouge, s. (eds): the jurassic of north-east greenland. geological survey of denmark and greenland bulletin 5, 31–49 (this volume). callomon, j.h. & birkelund, t. 1985: description of three new species. in: callomon, j.h.: the evolution of the jurassic ammonite family cardioceratidae. special papers in palaeontology 33, 78–86 (appendix). london: palaeontological association. carr, i.d. 1998: facies analysis and reservoir characterisation of jurassic sandstones from bjørnedal, central east greenland, 247 pp. unpublished ph.d. thesis, institute for sedimentology, university of reading, uk. dietl, g. & callomon, j.h. 1988: the orbis oolite in the upper bathonian (middle jurassic) of sengenthal/opf., franconian alb, and its significance for the correlation and subdivision of the orbis zone. stuttgarter beiträge zur naturkunde b 142, 31 pp. donovan, d.t. 1953: the jurassic and cretaceous stratigraphy and palaeontology of traill ø, east greenland. meddelelser om grønland 111(4), 150 pp. engkilde, m. & surlyk, f. 2003: shallow marine syn-rift sedimentation: middle jurassic pelion formation, jameson land, east greenland. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 813–863. fischer, p. 1880–1887: manuel de conchyliologie et de paléontologie, 1369 pp. paris: f. savy. haug, e. 1887: über die ‘polymorphidae’, eine neue ammonitenfamilie aus dem lias. neues jahrbuch für mineralogie, geologie und paläontologie 2, 89–163. larsen, m. & surlyk, f. 2003: shelf-edge delta and slope deposition in the upper callovian – middle oxfordian olympen formation, east greenland. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 931–948. madsen, v. 1904: on jurassic fossils from east greenland. meddelelser om grønland 29, 157–211. neumayr, m. 1875: die ammoniten der kreide und die systematik der ammonitiden. zeitschrift der deutschen geologischen gesellschaft 27, 854–892. neumayr, m. & uhlig, v. 1892: über die von k. abich im kaukasus gesammelten jurafossilien. denkschrift der akademie der wissenschaften in wien, mathematisch-naturwissenschaftliche klasse 59, 1–122. oppel, a. 1862: über jurassische cephalopoden. palaeontologische mitteilungen aus dem museum des koeniglich-bayerischen staates berlin 1(1–2), 1–125. page, k.n. 1989: a stratigraphical revision for the english lower callovian. proceedings of the geologists’ association (london) 100, 363–382. price, s.p. & whitham, a.g. 1997: exhumed hydrocarbon traps in east greenland: analogs for the lower–middle jurassic play of northwest europe. american association of petroleum geologists bulletin 81, 196–221. quenstedt, f.a. 1886–1887: die ammoniten des schwäbischen jura. ii. atlas, 55–90. stuttgart: schweizerbart. sokolov, d.n. & bodylevsky, v.i. 1931: juraund kreidefaunen von spitsbergen. skrifter om svalbard og ishavet 35, 151 pp. sowerby, j. 1813–1821: the mineral conchology of great britain 1–6, 1236 pp. london: j. de c. sowerby. spath, l.f. 1932: the invertebrate faunas of the bathonian– callovian deposits of jameson land (east greenland). meddelelser om grønland 87(7), 158 pp. surlyk, f. 1977: stratigraphy, tectonics and palaeogeography of the jurassic sediments of the areas north of kong oscars fjord, east greenland. bulletin grønlands geologiske undersøgelse 123, 56 pp. surlyk, f. 1978: submarine fan sedimentation along fault scarps on tilted fault blocks (jurassic–cretaceous boundary, east greenland). bulletin grønlands geologiske undersøgelse 128, 108 pp. surlyk, f. 1990: a jurassic sea-level curve for east greenland. palaeogeography, palaeoclimatology, palaeoecology 78, 71– 85. surlyk, f. 1991: sequence stratigraphy of the jurassic – lowermost cretaceous of east greenland. american association of petroleum geologists bulletin 75, 1468–1488. surlyk, f. 2003: the jurassic of east greenland: a sedimentary record of thermal subsidence, onset and culmination of rifting. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 659–722. surlyk, f., callomon, j.h., bromley, r.g. & birkelund, t. 1973: stratigraphy of the jurassic – lower cretaceous sediments of jameson land and scoresby land, east greenland. bulletin grønlands geologiske undersøgelse 105, 76 pp. surlyk, f., noe-nygaard. n. & dam, g. 1993: high and low resolution sequence stratigraphy in lithological prediction – examples from the mesozoic around the northern north atlantic. in: parker, j.r. (ed.): petroleum geology of northwest europe: proceedings of the 4th conference, 199–214. london: geological society. therkelsen, j. & surlyk, f. 2004: the fluviatile bristol elv formation, a new middle jurassic lithostratigraphic unit from traill ø, north-east greenland. in: stemmerik, l. & stouge, s. (eds): the jurassic of north-east greenland. geological survey of denmark and greenland bulletin 5, 19–29 (this volume). von siemiradzky, j. 1891: fauna kopalna warstw oksfordzkich i kimerydzkich w okregu krakowskim i przyleglych czésciach królestwa polskiego. czesc i: glowonigi. parnietnik akademii umietjosci w krakowie, wydzial matematycznoprzyrodniczy, tomu oseimnastego zeszyt i 18, 92 pp. vosgerau, h., alsen, p., carr, i.d., therkelsen, j., stemmerik, l. & surlyk, f. 2004: jurassic syn-rift sedimentation on a seawards-tilted fault block, traill ø, north-east greenland. in: stemmerik, l. & stouge, s. (eds): the jurassic of north-east greenland. geological survey of denmark and greenland bulletin 5, 9–18 (this volume). geus bulletin no 5.pmd 29-10-2004, 11:1438 39 plate 1 geus bulletin no 5.pmd 29-10-2004, 11:1439 40 plate 1 all specimens are figured natural size. fig. 1. kepplerites cf. tenuifasciculatus callomon. adult macroconch. mguh 25762 from ggu 429773. fig. 2. cadoceras sp. a: mguh 25763 from ggu 429772b. b: mguh 25764 from ggu 429772a. c: mguh 25765 from ggu 429772c. fig. 3. cadoceras sp. mguh 25766 from ggu 429772e. fig. 4. cadoceras cf. nordenskjoeldi callomon & birkelund a: mguh 25767 from ggu 429768a. b: mguh 25768 from ggu 429768c. c: mguh 25769 from ggu 429768b. fig. 5. cadoceras sp. indet. mguh 25770 from ggu 429767. geus bulletin no 5.pmd 29-10-2004, 11:1440 41 1 2 3 5 4 a c b a b c geus bulletin no 5.pmd 29-10-2004, 11:1441 42 plate 1appendix description of a new species of ammonite, kepplerites tenuifasciculatus n. sp., from the middle jurassic, lower callovian of east greenland john h. callomon a new species of ammonite, kepplerites tenuifasciculatus n. sp., is described. its type locality is at fossilbjerget in central jameson land, east greenland and its type horizon lies in the apertum zone, the lowest zone in the lower callovian stage of the middle jurassic. it has a narrow stratigraphical range and hence makes a good guide-fossil for stratigraphical time correlation. keywords: ammonite, east greenland, fossilbjerget, jameson land, kepplerites tenuifasciculatus n. sp., middle jurassic department of chemistry, university college london, 20 gordon street, london wc1h 0aj, uk. e-mail: johncallomon@lineone.net the revision of the biostratigraphy and biochronology of the ammonite faunas of the middle jurassic of east greenland has revealed a number of hitherto unknown species that characterize very narrow chronostratigraphic intervals and hence make excellent guide-fossils for time correlations (callomon 1993). one of these is now described. it represents a transient – chronospecies of some authors – in the evolution of one of the two evolutionary lineages, the kosmoceratidae, whose members inhabited east greenland during the late bathonian and callovian, the other being the much longer-ranging cardioceratidae. superfamily stephanocerataceae neumayr 1875 family kosmoceratidae haug 1887 genus kepplerites neumayr & uhlig 1892 type species. amm. keppleri oppel 1862. kepplerites tenuifasciculatus n. sp. plate 1, figs 1–2 1993 kepplerites sp. nov. j [tenuifasciculatus ms] callomon, p. 103. holotype. plate 1, fig. 1, mguh 25310 from ggu 185614a (t. birkelund and c. heinberg collection 1974) jameson land, fossilbjerget, section 43, bed 14, horizon j27 (figs 1, 2). other material. paratypes i, ii, mguh 25311–312 from ggu 185614b, c; paratype iii, mguh 25313 from jhc 4475 (plate 1, fig. 2; t. birkelund and j.h. callomon collection 1971), same locality and bed; jhc 4469–4471, section 42, bed 18 (see fig. 2); ggu 144191–192 from south slopes of mount mikael bjerg, section 31 (birkelund coll. 1971; fig. 1). numerous other specimens seen in situ were too poorly preserved to be worth collecting. stratigraphical horizon. the southern slopes of fossilbjerget are marked by three ridges, running southwards, on which sections have been recorded, numbered 41–43 from east to west (callomon 1993, fig. 1). geus bulletin no 5.pmd 29-10-2004, 11:1442 43 the sections span the shaly fossilbjerget formation underlain by the sandy pelion formation and capped by the sandstones and shales of the olympen formation (surlyk et al. 1973; larsen & surlyk 2003). the three sections differ only in detail. section 43 is shown in weathering-profile in figure 2. it was previously shown in outline by surlyk et al. (1973, fig. 23). lithologically, the sediments consist predominantly of shaly siltstones or very fine-grained sandstones, barely consolidated except in concretionary layers that punctuate the succession as markers or in scattered calcitic or phosphatic concretions. some of the beds are highly glauconitic, indicating condensed, sediment-starved intervals and comparison with adjacent areas, e.g. at the mountains olympen and mikael bjerg (fig. 1), shows that the succession incorporates numerous hiatuses of variable durations. the age-diagnostic ammonites can be recovered only from the hard beds. their biostratigraphy in terms of faunal horizons, however, is with only few exceptions as close to complete as present knowledge allows (callomon 1993, fig. 4). the only horizon in the interval under discussion not so far recognized in the fossilbjerget–olympen area is j23, that of kepplerites vardekloeftensis. it may have been lost in a hiatus, marked by a sharp lithological break, under the glauconitic ironstones of beds 9–13 (fig. 2, section 43). the type-horizon of kepplerites tenuifasciculatus, j27, is bed 14, an indurated, shaly, non-glauconitic, light brown fine-grained sandstone, 0.2 m thick, lying with sharp contact on the hard, ferruginous and highly glauconitic sandstone marker of bed 13 below (fig. 2, section 43). the contact probably marks another hiatus that would account for the considerable break, both in composition and morphologies, between the faunas of beds 13 and 14. the fauna of bed 13 consists predominantly of cadoceras (apertum), with only minor kepplerites (cf. traillensis). that of bed 14 is dominated by monospecific scoresby sund jameson land scoresby land sortehat 71ºn 22ºw23ºw24ºw25 50 km0 olympen mikael bjerg fossilbjerget pelion ugleelv katedralen n ei ll k lin te r h ur ry in le t sc hu ch er t d al kong oscar fjord fle ming fj ord ca rls be rg fjo rd g re en la nd fig. 1. sketch-map of the jameson land area with place names mentioned in the text. geus bulletin no 5.pmd 29-10-2004, 11:1443 44 o ly m pe n n or de ns kj oe ld i a pe rt um c al yx is hm ae va ri ab ile c ra no ce ph al . lo w er c al lo vi an u pp er b at ho ni anfo ss ilb je rg et pe lio n section 43 section 42 va rd ek lø ft 100 50 0 0 1 2 3 j34–35 j31? j30 j29 j27 j24–6 j22 j21 j18 33 25 20 15 10 5 m 26 j30 3 km j29 j28 j27 j26 j25 j24 terebratulid brachiopods belemnoteuthid cephalopods ammonites pectinid bivalves other bivalves thalassinoides burrows j22 j21 j20 j19b j19a j18 faunal horizons 24 22 20 18 17 15 13 11 9 8 7 31 29 27 24 22 20 18 14 9 7 5 1 6 5 fm m bed m bedgp fig. 2. diagrammatic sections in weathering profile through the fossilbjerget formation at its type locality on the southern slopes of fossilbjerget, central jameson land; section 43 is located 3 km west of section 42. note that the beds in the two sections are numbered independently – discussion of bed numbers in the text refers to those in section 43. lithostratigraphy at left, standard chronostratigraphy – substages and zones – at right. diagonal hatching: glauconitic. numbers j18–j35: the ammonite faunal horizons recognized in jameson land (see callomon 1993). the horizon of kepplerites tenuifasciculatus is j27. cranocephal., cranocephaloide. geus bulletin no 5.pmd 29-10-2004, 11:1444 45 kepplerites (tenuifasciculatus) with only occasional crushed, indeterminate cadoceras. description. all the available material consists of crushed internal moulds and measurements of dimensions (table 1) are of limited value. the figured specimens are typical; both are complete adult macroconchs with strongly uncoiling seams on the last whorl. the microconchs remain unknown. the ribbing is characteristically dense and fine, the primaries rising retroradially on the umbilical wall, then swinging in a strongly forwards-directed curve on the umbilical shoulder into accentuated prorsiradiate ribbing on the whorlside, dividing into fasciculate sheaves (tenuifasciculate) of secondaries at about a third flank-height, rising uncurved to the venter, persisting with little or no loss of strength to the simple, somewhat sinuous peristome. there is no evidence of the lateral accentuation of the primaries into tubercles seen in other species of kepplerites, especially in the younger ones and subsequently in the descendant, kosmoceras. inner whorls are not seen, so whether the earliest stages already have tabulate venters is not known. comparisons. the evolution of major morphological characters in the genus kepplerites, leading to kosmoceras in the middle callovian, was very gradual. differentiation of successive transients relies on relatively minor variations of size and ribbing, often perceptible only by the trained eye in assemblages of more than a single specimen in which the range of intraspecific variability can be assessed. changes were not continuously orthogenetic: characters could ‘progress’ and ‘regress’ with time largely independently, leading to frequent partial homoeomorphies. in descending order: 1. kepplerites traillensis donovan 1953 (plate 17, figs 1a, b, holotype; plate 18, figs 1a, b), faunal horizons 24–26, is similar in size, coiling and style of ribbing but significantly less densely ribbed, with only c. 31 primaries per whorl (before the onset of the modifications on the final stage of the adult bodychamber found in all the kosmoceratidae). the type material came from mount morris bjerg on traill ø, about 5 km north-east of the coast of kong oscar fjord to the south-west. it was found in isolation, both stratigraphically and faunistically, so the position of its faunal horizon in the general succession has to be deduced by correlation with the more continuous successions in jameson land. the closest resemblance is to the forms found also on the southern slopes of fossilbjerget, sections 42 and 43, horizons 24–26, as minor components in faunas dominated by cadoceras apertum (callomon & birkelund 1985). k. traillensis is also, among all the known faunas of greenland, the one closest to the type-species k. keppleri (oppel): cf. buckman (1922, plate 289a, b, lectotype, evolute inflated variant with slightly tabulate venter); quenstedt (1886, plate 77, figs 1–5, s. germany); page (1989, fig. 5.1a, b, england). the resemblance is close but may not be exact, so that both specific names are retained for the time being. it indicates, however, a close time correlation and provides the basis for the assignment of the apertum zone already to the lower callovian. 2. kepplerites vardekloeftensis callomon 1993 (p. 102), faunal horizon 23. the holotype (spath 1932, plate 25, figs 2a, b, complete adult) and paratype (spath 1932, plate 25, figs 1a, b, complete adult phragmocone) came from the calyx limestone, a prominent concretionary marker-bed in the fossilbjerget formation along the length of the outcrops above neill klinter, the line of cliffs on the west side of hurry inlet and traceable inland as far as katedralen on ugleelv (fig. 1); level 560 m in sections of rosenkrantz reproduced by spath (1932, p. 126, fig. 10). the species resembles k. tenuifasciculatus in coiling, size and density of primary ribbing but the secondary ribbing is coarser and fades on the table 1. measurement of dimensions holotype paratype iii maximum diameter 137 mm 140 mm septate to c. 100 mm c. 90 mm length of bodychamber, whorl 0.70 0.75 fractional umbilical width 0.23 0.22 at last septum fractional umbilical width c. 0.34 c. 0.33 at aperture primary ribs per whorl – at diameter 125–130 mm 51 57 – at diameter 85–95 mm 40 46 ratio of secondaries:primaries 3.8 3.4 around last septum geus bulletin no 5.pmd 29-10-2004, 11:1445 46 bodychamber. the two species are closely homoeomorphic but stratigraphically separated by the transients described above, which differ appreciably. 3. kepplerites svalbardensis sokolov & bodylevsky (1931, p. 79, plate 5, figs 1, 2) resembles k. tenuifasciculatus in coiling and finesse of ribbing, but is smaller: adult size 105 mm, septate to 70 mm. the ribbing (45 primaries per whorl) differs, however, on the whorl-side in that, after the initial forwards twist at the umbilical margin, it curves backwards again, the secondaries reaching the venter rectiradially. the species occurs in greenland as a minor component in fauna 22, the dominant element of which is k. peramplus spath. the latter is so distinctive, characterised by its great size (adult diameters 200 mm or more), involute and compressed inner whorls (see also dietl & callomon 1988, figs 4, 5), that there seems little doubt about the separate biospecific identities of the two taxa. they are not linked by intermediates and represent one of the rare cases in which the specific diversity at one horizon of an evolving generic clade rises above the monospecific. the faunal horizon can be followed from southern hurry inlet (types of k. peramplus) as far as the mountains mikael bjerg and fossilbjerget, sections 42 and 43 (see above; fig. 1). upwards, the record of kepplerites in greenland becomes tenuous. occasional specimens have been found in the nordenskjoeldi zone, horizons 28–29, but the preservation is too poor to be able to say much of interest other than that they are still of the general appearance and size of k. traillensis or k. tenuifasciculatus. the next horizon to yield keppleritids is horizon 32. the forms at this horizon are, however, quite distinct: small, evolute and round-whorled, typical of the chronosubgenus gowericeras that makes an abrupt appearance in much of northern europe and thereby characterizes the koenigi zone. age and distribution. lower callovian, apertum zone, faunal horizon 27. central jameson land, southern slopes of fossilbjerget, sections 42–43, and around mikael bjerg (fig. 1), sections 31, 33. geus bulletin no 5.pmd 29-10-2004, 11:1446 47 plate 1 geus bulletin no 5.pmd 29-10-2004, 11:1447 48 plate 1 complete adults, natural size; arrows mark the position of the last septum at the onset of the adult bodychamber. fig. 1. kepplerites tenuifasciculatus n. sp. holotype, mguh 25310 from ggu 185614a. fossilbjerget, section 43, bed 14, faunal horizon j27. lower callovian, apertum zone. fig. 2. kepplerites tenuifasciculatus n. sp. paratype iii, mguh 25313 from jhc 4475. section, bed and faunal horizon as above. geus bulletin no 5.pmd 29-10-2004, 11:1448 49 1 2 geus bulletin no 5.pmd 29-10-2004, 11:1449 geological survey of denmark and greenland bulletin 6, 57-66 57geological survey of denmark and greenland bulletin 6, 57–66 © geus, 2004 palaeoproterozoic age of a basement gneiss complex in the charcot land tectonic window, east greenland caledonides kristine thrane the charcot land tectonic window exposes crystalline basement gneisses, which form part of the foreland of the east greenland caledonides. these gneisses were previously believed to be archaean in age, on the basis of imprecise k-ar analyses carried out in the early 1980s on hornblende from amphibolitic bands and inconclusive rb/sr isotope data. new u-pb singlezircon ion microprobe analyses on the gneisses of the window yield upper intercept ages of 1916 ± 21 and 1928 ± 11 ma, and are interpreted to represent the age of crystallisation of the igneous protolith. the foreland gneisses of the charcot land window are similar in age to parts of the allochthonous gneiss complexes of structurally overlying thrust sheets, but the two terranes have different lithological and structural characteristics. no archaean rocks have been identified with certainty in any of the east greenland caledonian foreland windows. keywords: caledonides, east greenland, geochronology, palaeoproterozoic geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. present address: geological institute, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. email: kthrane@geol.ku.dk the charcot land window is located between latitudes 71°45′n and 72°15′n in the south-western part of the 1300 km long east greenland caledonides (fig. 1). it is one of a series of tectonic windows, interpreted to expose parautochthonous foreland (higgins & leslie 2000), found along the western border of the caledonian orogen. the gåseland window is exposed to the south at c. 70°n (wenk 1961; phillips et al. 1973), and the målebjerg and eleonore sø windows to the north between latitudes 73°30′–74°15′n (leslie & higgins 1998; higgins & leslie 2000, 2004, this volume). further to thenorth, the foreland basement is also exposed in western dronning louise land (76°–77°n; e.g. strachan et al. 1992), and in the small nørreland window (78°40′n; hull & friderichsen 1995). apart from these restricted areas the foreland of the east greenland caledonides is not well known, as it is largely concealed beneath the inland ice to the west. the charcot land window (fig. 2) exposes a basement complex of grey gneisses with amphibolite bands, overlain by a cover of low to medium-grade supracrustal rocks including carbonates, volcanic rocks and siliciclastic sediments (charcot land supracrustal sequence; steck 1971). both basement and cover are cut by major palaeoproterozoic granite intrusions (hansen et al. 1981). structurally overlying the charcot land window are caledonian thrust sheets comprising crystalline gneiss complexes (flyverfjord infracrustal complex) and overlying high-grade metasedimentary rocks (krummedal supracrustal sequence; henriksen & higgins 1969, 1976). the crystalline gneisses making up the flyverfjord infracrustal complex in hinks land, south-east of charcot land (fig. 2), are of late archaean age (c. 2700 ma; rex & gledhill 1974; thrane 2002). the flyverfjord infracrustal complex extends both southwards and northwards (fig. 1); however, north of 72°50′n very similar looking orthogneisses have yielded palgeus bulletin 6.pmd 10-02-2005, 09:5457 58 ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ c a le d o n ia n f o ld b el t in la n d ic e palaeogene basalts palaeogene intrusions wandel sea basin: carboniferous–palaeogene east greenland basin: carboniferous–cretaceous sediments devonian – continental sediments crystalline basement thrust fault/shear zone post-caledonian late to post-caledonian caledonian fold belt caledonian foreland 70° 74° 78° 35° wandel hav 100 km jameson land scoresby sund 25° traill ø bessel fjord kejser franz joseph fjord kong oscar fjord dronning louise land fig. 2 danmarkshavn w es te rn li m it of de fo rm at io n ▲ ▲ peary land k ro np rin s c hr ist ian la nd caledonian granites neoproterozoic–ordovician sediments (east greenland) palaeo-mesoproterozoic sediments and basalts (north greenland) neoproterozoic–silurian sediments (north greenland) neoproterozoic–silurian sediments (north greenland) crystalline gneisses and volcano-sedimentary complexes palaeo-mesoproterozoic sediments and basalts (north and north-east greenland) high-grade metasediments (krummedal supracrustal sequence) charcot land window gåseland window eleonore sø window målebjerg window 25° 15° 15° 35° 82° ▲ ▲ ▲ fig. 1. geological map of the east greennland caledonides. frame indicates study area shown at larger scale in fig. 2. geus bulletin 6.pmd 10-02-2005, 09:5458 59 aeoproterozoic ages (c. 1900 ma; rex & gledhill 1981; kalsbeek et al. 1993; thrane 2002). the krummedal supracrustal sequence, that overlies the flyverfjord infracrustal complex as well as the similar gneiss complexes between latitudes 70° and 76°n, was deposited later than 1100 ma ago, and underwent high grade metamorphism with generation of granites c. 940 ma ago (kalsbeek et al. 2000). the contact between the krummedal supracrustal sequence and the structurally overlying neoproterozoic–ordovician sedimentary succession is a shear zone, interpreted as an extensional detachment or locally a thrust (escher & jones 1998; leslie & higgins 1998). the western border of the main neoproterozoic–ordovician outcrop in the fjord region from 72°–75°n is a late orogenic extensional fault system (e.g. hartz & andresen 1995). the c. 13 km thick neoproterozoic eleonore bay supergroup succession (sønderholm & tirsgaard 1993) together with the vendian tillite group (800–1000 m) and cambro-ordovician shelf sequence (up to 4 km), constitute the highest part of the uppermost caledonian thrust sheet (higgins et al. 2004). in 1968, the southern part of the charcot land area was mapped by the geological survey of greenland (henriksen & higgins 1969, 1976; steck 1971), during a systematic regional mapping project in the scoresby sund region (70°–72°n). during the 1997 and 1998 survey expeditions to the kong oscar fjord region (72°–75°n), the northern part of charcot land was mapped by friderichsen & thrane (1998). sample collections were made for ion microprobe zircon studies, the results of which are reported here. geological setting the foreland rock units exposed within the charcot land window include crystalline basement orthogneisses with interleaved amphibolite bands, that are overlain by a c. 2000 m thick succession of various metasedimentary rocks (including white marble) and metamorphosed basic extrusive and intrusive rocks (charcot land supracrustal sequence). these units are cut by two major granitoid intrusions which crop out widely in central charcot land (fig. 2); one is a hornblende-biotite quartz diorite to granodiorite body and the other a pegmatitic muscovite granite. in southernmost charcot land a hornblende gabbro body intrudes the basement gneisses. tillit nunatak in south-west charcot land takes its name from a sheared tillite resting unconformably on the granodiorite intrusion. this tillite has been correlated with the vendian tillites of the fjord zone (moncrieff 1989), as has a similar tillite in the gåseland window. the charcot land tillite is very little deformed but does contain a planar cleavage. the charcot land supracrustal sequence of southernmost charcot land is characterised by greenschist facies metamorphism. metamorphic grade increases towards the north-east and reaches upper amphibolite facies at ~ 72°n latitude (steck 1971). further to the north, the metamorphic grade again decreases (friderichsen & thrane 1998). the eastern and western boundaries of the charcot land window are marked by major caledonian thrusts (fig. 2). the western thrust dips gently to the t t overlying thrust units tillit nunatak formation metasediments marble thrust fold trendmetavolcanics hornblende gabbro 30˚ 72˚ 30˚ charcot land c ha rc ot l an d se qu en ce an tillit nunatak daugaard-jensen glets cher nvf hinks land 426040 1928 ± 11 ma 29˚ muscovite granite quartz diorite basement gneiss and amphibolite 426041 1916 ± 21 ma 10 km 29˚ fig. 2. geological map of the charcot land window, showing major lithological divisions and sample locations. modified from henriksen & higgins (1976). the rock types of the overlying thrust sheets are undifferentiated, and include gneisses of the flyverfjord infracrustal complex and the krummedal supracrustal sequence. an, alfabet nunatakker; nvf, nordvestfjord. geus bulletin 6.pmd 10-02-2005, 09:5459 60 west and is only exposed on the alfabet nunatakker, while the eastern thrust dips to the east and can be traced for more than 25 km in western hinks land (higgins 1982). north-east of charcot land, on the north-east side of innermost nordvestfjord, a very large-scale, recumbent, west-vergent isoclinal anticline makes up the eastern boundary of the window (fig. 3). this fold nappe is sitting in the hanging wall and is made up by the cover rocks from outside of the window. the nappe, first recognisedbyeduardwenkandhelgebacklund in 1934 (backlund in koch 1955; wenk 1956), is outlined by a conspicuous white marble unit that has an isoclinal fold closure some 25 km north-west of its first appearance opposite the front of daugaard-jensen gletscher. however, the narrow closure that wenk and backlund observed from fjord level accounts for less than half of the magnitude of this fold nappe. over its entire length, the structure could easily accommodate a translative movement of the proportions necessary to form the charcot land window, a minimum of 40 km displacement (friderichsen & thrane 1998). structures the broad structure of the charcot land window is an elongate n–s-trending dome (fig. 2). within the window, foliation trends have a general nw–se strike and dip at shallow angles to the east. lineations plunge at moderate angles eastwards, between north-east and south-east. major folds have trends between e–w and nw–se, while minor folds have more variable trends and plunge directions. the crystalline basement terrain east and north-east of charcot land, that forms part of the structurally overlying thrust sheet, exhibits a different pattern of structures. friderichsen & thrane (1998) mapped a broad area of the allochthonous gneisses north-east of charcot land. they record the general foliation trends in eastern parts of their area as ne–sw to n–s with an eastwards dip, with linear elements plunging at moderate to shallow angles to the east or southeast; these were interpreted as caledonian structures. the presence of major, recumbent, nw-verging folds suggests the ne–sw trends might be a response to caledonian compression, while the n–s trends may relate to caledonian extension, mainly seen as ductile top-down-to-the-east faults. in the western part of their area, friderichsen & thrane (1998) noted the structural pattern as more diffuse, perhaps due to a fig. 3. part of the major, recumbent, west-verging isoclinal fold that defines the eastern boundary of the charcot land window (hanging wall). the isocline is outlined by the two white marble bands (m) which form a closure just beyond the left side of the photograph. the black units are amphibolite. north-east of innermost nordvestfjord (see fig. 2); profile height about 1000 m. geus bulletin 6.pmd 10-02-2005, 09:5460 61 less intense caledonian overprinting of older and more complex pre-caledonian structures. in general the foliation trends are ne–sw, with nw-plunging lineations. although the general foliation trends within the charcot land window are also broadly nw–se, like those in the eastern areas of the overlying thrust sheets, the structures in charcot land are not interpreted as caledonian. lineations in charcot land also have much steeper eastward dips than in the thrust sheets. further evidence that the caledonian overprint within the window was generally weak is that the large palaeoproterozoic granite intrusions cutting the basement gneisses in charcot land show no evidence of deformation. the only certain caledonian structures are the planar cleavage in the diamictites of the vendian tillit nunatak formation. the major thrust along the east side of the charcot land window is well-defined in north-west hinks land (fig. 2), but its northward continuation on the northeast side of nordvestfjord was previously uncertain (henriksen & higgins 1976). on the basis of their 1997 studies, friderichsen & thrane (1998) proposed that the movement might have been taken up by the major nappe-like fold on the north-east side of the charcot land window (fig. 3). the trend and westward sense of overturning of this nappe-like fold support the viewpoint that it is a caledonian structure, although no other comparable structures were observed within the window. a thick low-grade succession of dark phyllitic shales observed in the crest of the fold has also been traced in the western nunataks of charcot land, substantially increasing the possible extent of the structure. in addition a thrust associated with this nappe-like major fold suggests that the total displacement was probably significantly more than the 40 km proposed by henriksen & higgins (1976) and henriksen (1986). previous geochronological studies hansen et al. (1981) analysed hornblende from amphibolite bands within the charcot land basement gneisses, and obtained k-ar ages of 2097 ± 105 ma and 2855 ± 145 ma. these were regarded as minimum ages for the amphibolite units of the crystalline complex. rb-sr whole rock analyses on samples of the gneisses were undertaken by hansen (1976), but while the results did not yield any exact ages, there were indications that the gneisses might be more than 2200 ma old. based on lithology and age an archaean origin comparable to that of the flyverfjord basement gneisses was thought likely. rb-sr whole rock analyses were also undertaken on samples of the widespread pegmatitic muscovite granite body, that intrudes both the basement gneisses and the supracrustal rocks in charcot land (hansen et al. 1981); an age of c. 1850 ma was considered to indicate the approximate time of intrusion. k-ar analyses on large muscovite crystals from the same granite yielded ages of 1760 ± 60 and 1870 ± 60 ma, interpreted as a minimum age of emplacement (hansen et al. 1981). based on field observations, steck (1971) argued that the granite was intruded after the main metamorphism of the supracrustal rocks, which was suggested to have taken place about 1900–1850 ma ago (hansen et al. 1981). a young rb-sr biotite age of 402 ± 10 ma from an amphibolite, together with a biotite-feldspar-wholerock isochron age of 402 ± 8 ma for samples of the pegmatitic granite, were considered to reflect a caledonian greenschist facies metamorphic overprint (hansen et al. 1981). isotopic data from orthogneiss units in some of the other foreland areas exposed along the margin of the caledonian orogen in east greenland have yielded palaeoproterozoic protolith ages, e.g. kalsbeek et al. (1993) and tucker et al. (1993) obtained palaeoproterozoic ages for basement gneisses from the western part of dronning louise land (77°n, 25°w) and quartz porphyry bodies and grey granites within the eleonore sø window have yielded similar ages (higgins & leslie 2004, this volume). none of the so far dated foreland rock units have yielded convincing archaean ages. geochronology samples two representative samples from the basement gneiss complex were studied. sample 426040 is a fine-grained, light grey, granodioritic orthogneiss, and sample 426041 a dark grey, coarse-grained, granodioritic orthogneiss (fig. 4). both rock units exhibit a penetrative gneissosity and mineral lineation, and are cut by deformed light-coloured granite veins and undeformed pegmatitic muscovite granite veins. in the field, the age relationships between the two orthogneisses are very clear; the fine-grained orthogneiss (426040) intrudesandcuts thecoarse-grainedorthogneiss (426041). the limited field work undertaken in charcot land in 1997, and that undertaken earlier during the 1968– geus bulletin 6.pmd 10-02-2005, 09:5461 62 1972 scoresby sund expeditions (henriksen 1986), indicate both gneiss types to be part of the regional basement gneiss complex, the oldest rocks exposed in charcot land. analytical methods for the ion microprobe study u-pb dating of zircons was undertaken using the cameca ims 1270 ion probe at the nordsim laboratory, swedish museum of natural history, stockholm. approximately 50 zircon grains from each sample were hand-picked, and mounted in a transparent epoxy resin together with reference zircons 91500 (from ontario, canada, with a weighted average 207pb/206pb age of 1065 ma; wiedenbeck et al. 1995). the zircon grains were polished sufficiently to expose any potentially older cores. the mounts were examined by reflected light microscopy and by backscatter imaging ina scanningelectron microscope and then coated with c. 30 nm of gold. analytical procedures are similar to those described by schuhmacher et al. (1994) and whitehouse et al. (1997). calibration of pb/u ratios follows procedures similar to those used by the ion probegroup at the australian national university (williams 1998), and is based on observed relationships between pb/u and uo 2 /u, during the same analytical run. results are given as discordia line intercept ages. the ages were calculated using isoplot/ex (ludwig 1999). analytical results the u-pb zircon data (table 1) are presented in conventional concordia diagrams (fig. 6; 1 f error ellipses). sample 426041 most of the zircons in sample 426041 are elongate, prismatic and clear, with sizes ranging from 100 to 400 mm, most commonly between 200 and 250 mm. they contain solid homogenous cores, some of which show oscillatory zonation. the rims are broad, metamict and show zonation. from the backscatter images, it is not clear whether the rims are metamorphic or simply alteration rims (fig. 5). however, there is a clear chemical variation, with the rims having much lower th/u ratios (0.02–0.07) than the cores (0.19– 0.36), and also containing more common lead (table 1). the variation in the th/u ratios indicates that the rims may have a metamorphic origin, while the higher common lead content may be a result of alteration of the rims. in the concordia diagram (fig. 6a), all the analyses (rims as well as cores) fall on the same discordia line. the upper intercept age is 1916 ± 21 ma and the lower intercept date of 443 ± 25 ma (mswd = 10.9), where the upper intercept is interpreted as the best approximation of the crystallisation age of the gneiss protolith. most of the rims plot on the discordia line, indicating that they must have a palaeoproterozoic origin very close in age to the protolith, e.g. the rock must have been exposed to metamorfig. 4. dark grey orthogneiss (426041) cut by a broad dyke of light grey orthogneiss (426040). both are cut by later pegmatitic granite veins. water flowing from right to left produces the colour differences between wet and dry rocks. fig. 5. backscatter image of zircon number 7 from sample 426041. the circles indicate the ion probe analysis sites. geus bulletin 6.pmd 10-02-2005, 09:5462 63 table 1. sims u-th-pb analytical data and derived ages sample/ u pb th th/u f % 207pb ±σ 207pb 206pb ±σ disc. ages (ma) spot# ppm ppm ppm meas. common 206pb % 235u % 238u % % 207pb ±σ 207pb ± 206pb ±σ 206pb 235u 238u 426040 9c 351 137 86 0.25 0.08 0.1170 0.29 5.275 0.82 0.327 0.77 –5.3 1911 5 1865 7 1824 12 6c 473 148 134 0.28 0.07 0.1125 0.33 4.041 0.82 0.261 0.75 –21.1 1840 6 1642 7 1493 10 5c 717 192 279 0.39 1.10 0.1034 0.46 2.713 0.92 0.190 0.79 –36.3 1686 8 1332 7 1123 8 1c 1583 245 727 0.46 1.29 0.0998 0.55 1.739 1.00 0.126 0.83 –55.8 1620 10 1023 6 767 6 3c 628 122 190 0.30 1.14 0.0996 0.92 2.243 1.26 0.163 0.87 –42.7 1616 17 1195 9 975 8 4c 856 158 362 0.42 0.18 0.0971 0.42 2.062 1.14 0.154 1.06 –44.1 1569 8 1136 8 923 9 2c 752 134 194 0.26 1.49 0.0958 0.59 2.007 1.12 0.152 0.96 –43.8 1543 11 1118 8 912 8 7r 7482 463 250 0.03 4.80 0.0527 5.96 0.425 6.01 0.058 0.77 15.9 317 130 359 18 366 3 426041 2c 302 110 83 0.27 0.09 0.1165 0.40 4.883 0.93 0304 0.83 –11.6 1904 7 1799 8 1711 13 13c 247 93 88 0.36 0.11 0.1143 0.49 4.857 1.19 0.308 1.09 –8.3 1869 9 1795 10 1732 17 14c 503 138 94 0.19 0.40 0.1115 0.42 3.635 0.93 0.236 0.83 –27.7 1824 8 1557 7 1368 10 4c 364 107 123 0.34 0.21 0.1107 0.41 3.707 0.97 0.243 0.87 –25.1 1810 7 1573 8 1402 11 2r 475 127 106 0.22 0.24 0.1086 0.39 3.424 0.95 0.229 0.87 –27.9 1775 7 1510 8 1328 10 7c 780 212 162 0.21 0.16 0.1067 0.25 3.415 0.96 0.232 0.93 –25.3 1745 5 1508 8 1345 11 9c 618 136 175 0.28 0.53 0.1057 0.45 2.711 1.15 0.186 1.06 –39.4 1727 8 1332 9 1100 11 1c 683 134 192 0.28 0.36 0.0999 0.41 2.288 0.93 0.166 0.84 –41.9 1621 8 1208 7 991 8 7r 3705 594 95 0.03 0.40 0.0935 0.29 1.827 0.88 0.142 0.83 –45.8 1498 5 1055 6 854 7 8r 3445 514 73 0.02 0.83 0.0924 0.33 1.727 0.90 0.136 0.83 –47.3 1475 6 1019 6 820 6 9r 2780 289 132 0.05 0.69 0.0750 0.58 0.992 1.09 0.096 0.92 –46.7 1068 12 700 6 591 5 15r 3761 349 159 0.04 1.59 0.0732 0.63 0.867 1.05 0.086 0.84 –49.8 1019 13 634 5 531 4 3r 3824 354 93 0.02 0.80 0.0693 0.52 0.823 0.98 0.086 0.83 –43.2 909 11 609 5 532 4 11r 3622 294 108 0.03 1.67 0.0598 0.80 0.630 1.16 0.076 0.85 –21.2 597 17 496 5 475 4 5r 3622 280 255 0.07 2.27 0.0548 0.87 0.542 1.22 0.072 0.86 –11.4 403 19 440 4 447 4 errors on ratios and ages are quoted at 1 σ level. f % common is the fraction of common 206pb estimated from the measured 204pb. disc. % refers to the degree of discordance of the zircon analysis. σ u pb th % σ σ σ σ σ 1911 5 1865 7 1824 12 1840 6 1642 7 1493 10 1686 8 1332 7 1123 8 1620 10 1023 6 767 6 1616 17 1195 9 975 8 1569 8 1136 8 923 9 1543 11 1118 8 912 8 317 130 359 18 366 3 1904 7 1799 8 1711 13 1869 9 1795 10 1732 17 1824 8 1557 7 1368 10 1810 7 1573 8 1402 11 1775 7 1510 8 1328 10 1745 5 1508 8 1345 11 1727 8 1332 9 1100 11 1621 8 1208 7 991 8 1498 5 1055 6 854 7 1475 6 1019 6 820 6 1068 12 700 6 591 5 1019 13 634 5 531 4 909 11 609 5 532 4 597 17 496 5 475 4 403 19 440 4 447 4 g e u s b ulletin 6.pm d 10-02-2005, 09:54 63 64 phism soon after crystallisation of the protolith, and these rims have later suffered pb loss. only a few rims cluster around the lower intercept and could reflect the time when the gneiss suffered caledonian metamorphism. sample 426040 the morphologies of the zircons in 426040 are very similar to those in sample 426041. most have long prismatic shapes, but stubby grains also occur. the zircons are clear and range in size from 100 to 250 mm. both homogenous and oscillatory zoned cores are present, while the rims are often metamict. in this study mostly cores were analysed. they plot on a discordia line (fig. 6b) yielding intercept ages of 1928 ± 11 ma and 467 ± 18 ma (mswd = 0.74). two rims were analysed; these do not plot on the discordia line, but one of them plots on the concordia line, yielding a 206pb/ 238u age of 366 ± 3 ma. discussion the two analysed samples of basement gneisses from the charcot land window yield identical upper intercept ages for the protolith within error, 1928 ± 11 and 1916 ± 21 ma. the fact that the two orthogneisses yield the same age demonstrates that several distinct magma pulses were emplaced during this time period. if the rb-sr whole-rock age of c. 1850 ma reported by hansen et al. (1981) for the undeformed muscovite granite is valid, then the time of deformation and gneissification is bracketed between c. 1920 and 1850 ma. the palaeoproterozoic metamorphic effect seen in the growth of rims on zircons in sample 426041 might have been caused by the palaeoproterozoic intrusive activity. the lower concordia intercepts suggesting caledonian ages are indicative of some pb loss and alteration of the zircons (fig. 6), and can be linked to burial of the foreland beneath a several kilometre thick pile of caledonian thrust sheets. within the charcot land window limited caledonian deformation is recorded, e.g. development of a planar cleavage in the diamictites of the vendian tillit nunatak formation, and associated folding in the underlying basement gneisses (moncrieff 1989). however, caledonian deformation was not pervasive and, as noted above, the palaeoproterozoic muscovite granite body has no internal fabric. the new age determinations reported here indicate that the protolith of the basement gneiss complex in charcot land is palaeoproterozoic, rather than archaean as suggested by hansen (1976) and hansen et al. (1981). at the current state of knowledge there is, in fact, no convincing isotopic evidence of the presence of archaean basement rocks in any of the foreland windows along the western margin of the caledonian orogen. volcano-sedimentary sequences of palaeoproterozoic age are known in the charcot land and eleonore sø windows (higgins et al. 2001), but are not known in the allochthonous thrust sheets that overlie the foreland windows. similarly, the late mesoproterozoic – early neoproterozoic krummedal supracrustal sequence that is widely represented in the thrust sheets 0.4 0.3 0.2 0.1 0.4 0.3 0.2 0.1 0.0 0 1 2 3 4 5 6 7 400 800 1200 1600 2000a 426041 grey orthogneiss cut by 426040 8 9r 5 5 34 7r 1 9c 2r 1511 3 14 13 7c 4 2c intercepts at 443 + 25 & 1916 + 21 ma (mswd = 10.9, n = 15) 20 6 pb / 23 8 u 0.0 0 1 2 3 4 5 6 7 400 800 1200 1600 2000b 426040 grey orthogneiss intruding 426041 intercepts at 467 + 18 & 1928 + 11 ma (mswd = 0.74, n = 6) 20 6 pb / 23 8 u 207pb / 235u 2 1 7 6 9 fig. 6. u-pb concordia diagrams for the two samples of orthogneiss from charcot land. a: sample 426041. b: sample 426040. error ellipses show 1 σ errors and errors on the intercept ages are 2 σ. geus bulletin 6.pmd 10-02-2005, 09:5464 65 structurally above these windows is not present within them. it could, therefore, be argued that the known archaean rock units, such as the allochthonous flyverfjord infracrustal complex south of latitude 72°50′n, represent a terrain that was accreted onto the palaeoproterozoic foreland. there remain many uncertainties in making regional assessments. for example, modern isotopic age data are still lacking for the crystalline basement rocks of the southern gåseland window, while metasedimentary rocks found as infolded layers within the allochthonous crystalline basement gneisses north of latitude 72°n (friderichsen & thrane 1998; thrane & friderichsen 1999) are of uncertain age. the assumption that these infolded metasedimentary layers are related to the thick developments of the late mesoproterozoic – early neoproterozoic krummedal sequence is not proven, and it cannot be excluded that they are equivalents of the palaeoproterozoic charcot land supracrustal sequence. the absence of representatives of the krummedal supracrustal sequence in the foreland windows could be considered as support for the view that this sequence was deposited far away from the flyverfjord crystalline basement and that these units were juxtaposed during the caledonian orogeny (watt & thrane 2001). the ages obtained from the basement gneiss complex in the charcot land foreland window are not dissimilar from the ages obtained on the allochthonous crystalline basement complexes north of 72°50′n (rex & gledhill 1981; kalsbeek et al. 1993; thrane 2002), although the lithological make-up of the two basement complexes is different. characteristic features of charcot land include the spectacular palaeoproterozoic granitoid intrusions, which occur both as netveining dykes and as major undeformed plutons. such characteristic features are uncommon in the allochthonous palaeoproterozoic crystalline basement rocks to the north. it is the author’s opinion that the grey orthogneisses in the allochthonous crystalline basement north of 72°50′n originally had much the same appearance as the charcot land orthogneisses, but that the former suffered much more intense deformation during the caledonian orogeny. thus, the charcot land basement gneisses and the allochthonous basement gneiss complexes north of 72°50′n could once have been parts of the same terrain prior to separation by caledonian thrusting, with the major nappelike structure on the north-east side of the charcot land window taking up the thrust movement. alternatively, the two palaeoproterozoic basement terrains may originally have been unrelated, and with different histories, but have been brought into close proximity by thrusting during the caledonian orogeny. acknowledgements the isotope data described in this paper were acquired at the nordsim laboratory at the swedish museum of natural history, stockholm. critical comments on the manuscript by johan d. friderichsen, adam a. garde, a.k. higgins and feiko kalsbeek are greatly appreciated. this project was based on funding from the danish natural science research council. david gee and clark friend are thanked for constructive reviews of the manuscript. references escher, j.c. & jones, k.a. 1998: caledonian thrusting and extension in frænkel land, east greenland (73°–73°30′n): preliminary results. in: higgins, a.k. & frederiksen, k.s. 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(ed.): express report: eastern north greenland and north-east greenland 1995, 11–21. unpublished report, geological survey of greenland, copenhagen. kalsbeek, f., nutman, a.p. & taylor, p.n. 1993: palaeoproterozoic basement province in the caledonian fold belt of northeast greenland. precambrian research 63, 163–178. kalsbeek, f., thrane, k., nutman, a.p. & jepsen, h.f. 2000: late mesoproterozoic to early neoproterozoic history of the east greenland caledonides: evidence for grenvillian orogenesis? journal of the geological society (london) 157, 1215–1225. koch, l. 1955: report on the expeditions to central east greenland 1926–1939. conducted by lauge koch. part ii. meddelelser om grønland 143(2), 642 pp. leslie, a.g. & higgins, a.k. 1998: on the caledonian geology of andreé land, eleonore sø and adjacent nunataks (73°30′– 74°n), east greenland. in: higgins, a.k. & frederiksen, k.s. (eds): caledonian geology of the east greenland caledonides 72°–74°n: preliminary reports from the 1997 expedition. danmarks og grønlands geologiske undersøgelse rapport 1998/28, 11–27. ludwig, k.r. 1999: isoplot/ex version 2.00 – a geochronological toolkit for microsoft excel. berkeley geochronology center, special publication 2. moncrieff, a.c.m. 1989: the tillite group and related rocks of east greenland: implications for late proterozoic palaeogeography. in: gayer, r.a. (ed.): the caledonide geology of scandinavia, 285–297. london: graham & trotman. phillips, w.e.a., stillman, c.j., friderichsen, j.d. & jemelin, l. 1973: preliminary results of mapping in the western gneiss and schist zone around vestfjord and inner gåsefjord, southwest scoresby sund. rapport grønlands geologiske undersøgelse 58, 17–32. rex, d.c. & gledhill, a. 1974: reconnaissance geochronology of the infracrustal rocks of flyverfjord, scoresby sund, east greenland. bulletin geological society of denmark 23, 49–54. rex, d.c. & gledhill, a. 1981: isotopic studies in the east greenland caledonides (72°–74°n) – precambrian and caledonian ages. rapport grønlands geologiske undersøgelse 104, 47– 72. schuhmacher, m., de chambost, e., mckeegan, k.d., harrison, t.m. & migeon, h. 1994: in situ dating of zircon with the cameca ims 1270. in: benninghoven, a. (ed.): secondary ion mass spectrometry sims ix, 919–922. chichester: john wiley & sons. sønderholm, m. & tirsgaard, h. 1993: lithostratigraphic framework of the upper proterozoic eleonore bay supergroup of east and north-east greenland. bulletin grønlands geologiske undersøgelse 167, 38 pp. steck, a. 1971: kaledonische metamorphose der praekambrischen charcot land serie, scoresby sund, ost-grönland. bulletin grønlands geologiske undersøgelse 97, 69 pp. (also meddelelser om grønland 192(3)). strachan, r.a., holdsworth, r.e., friderichsen, j.d. & jepsen, h.f. 1992: regional caledonian structure within an oblique convergence zone, dronning louise land, ne greenland. journal of the geological society (london) 149, 359–371. thrane, k. 2002: relationships between archaean and palaeoproterozoic basement complexes in the southern part of the east greenland caledonides: an ion microprobe study. precambrian research 113, 19–42. thrane, k. & friderichsen, j.d. 1999: the häsi bjerge orthogneiss complex, the foldebjerg orthogneisses and a brief account of a visit to the nordvestfjord nappe, east greenland caledonides. in: higgins, a.k. & frederiksen, k.s. (eds): geology of east greenland 72°–75°n, mainly caledonian: preliminary reports from the 1998 expedition. danmarks og grønlands geologiske undersøgelse rapport 1999/19, 37–41. tucker, r.d., dallmeyer, r.d. & strachan, r.a. 1993: age and tectonothermal record of laurentian basement, caledonides of ne greenland. journal of the geological society (london) 150, 371–379. watt, g.r. & thrane, k. 2001: early neoproterozoic events in east greenland. precambrian research 110, 165–184. wenk, e. 1956: alpines und ostgrönländisch-kaledonisches kristalline, eintektonish-petrogenetischervergleich. verhandlungen des naturforscher gesellschaftes in basel 67, 75–102. wenk, e. 1961: on the crystalline basement and the basal part of the pre-cambrian eleonore bay group in the south-west part of scoresby sund. meddelelser om grønland 168(1), 54 pp. whitehouse, m.j., claesson, s., sunde, t. & vestin, j. 1997: ion microprobe u-pb zircon geochronology and correlation of archaean gneisses from the lewisian complex of gruinard, north-western scotland. geochimica et cosmochimica acta 61, 4429–4438 wiedenbeck, m., allé, p., corfu, f., griffin, w.l., meier, m., oberli, f., von quardt, a., roddick, j.c. & spiegel, w. 1995: three natural zircon standards for u-th-pb, lu-hf, trace element and ree analyses. geostandards newsletter 19, 1–23. williams, i.s. 1998: u-th-pb geochronology by ion microprobe. in:mckibben, m.a., shanks iii, w.c. & ridley, w.i. (eds): applications of microanalytical techniques to understanding mineralizing processes. reviews in economic geology 7, 1–35. geus bulletin 6.pmd 10-02-2005, 09:5466 geological survey of denmark and greenland bulletin 20, 2010, 31–34 31 following the proposal of the offshore anholt wind-farm project with an energy capacity of 400 megawatt in the kattegat, southern scandinavia, an evaluation of the geotechnical properties of the subsurface of the area is required. as a first step to map the seabed geology the geological survey of denmark and greenland (geus) conducted a geophysical survey (leth et al. 2009) which, together with cone penetration tests and data from boreholes, lead to a greater understanding of the geological architecture and development of the 144 km2 survey area (figs 1, 2). methods we used a multibeam echo-sounder for detailed mapping of the bathymetry, and shallow seismic equipment and coring to map the shallow seabed geology including the distribution and thickness of the main geological units. a combination of two seismic devices (chirp and sparker systems) was chosen to ensure good penetration and high resolution. the sparker system provides data from the sea floor down to about 45 m into the seabed with a vertical seismic resolution in the order of 50 cm, while the chirp system provides high-resolution seismic data from the upper 5–10 m of the seabed with a vertical seismic resolution at decimetre scale. a side-scan sonar was used for mapping of surface sediments. at seven sites, boreholes were made to a depth of 40 m and selected intervals sampled. cone penetration tests were carried out at the same sites (fig. 2) and surface sediment samples collected for biological studies. geological setting and seabed the kattegat region is located in the transition zone between the fennoscandian shield and the danish basin (fig. 1), and studies of the pre-quaternary surface morphology show that the nw–se-trending anticlinorium follows the trend of late quaternary geology of a potential wind-farm area in the kattegat, southern scandinavia jørgen o. leth and bernhard novak bh-09 bh-01 bh-05 bh-02 56°33´n 11°15´e 2 km 14.5–15.0 15.0–15.5 15.5–16.0 16.0–16.5 16.5–17.0 17.0–17.5 17.5–18.0 18.0–18.5 18.5–19.0 19.0–19.5 depth (m) fig. 2. bathymetry of the survey area. the stars show sites with combined coring and cone penetration tests. the dashed line shows the position of the profile in fig. 3. sweden jylland sjælland kattegat precambrian and palaeozoic lower cretaceous upper cretaceous triassic tertiary jurassic 10°e 57°30´ 56°n 13°e 50 km ? study area lysegrund fennoscandian shield danish basin fig. 1. pre-quaternary geology of the kattegat. redrawn and simplified from lykke-andersen et al. (1993). © geus, 2010. geological survey of denmark and greenland bulletin 20, 31–34. open access: www.geus.dk/publications/bull 3232 dextral wrench faults that have repeatedly affected the fennoscandian border zone (lykke-andersen et al. 1993). the latest major tectonic event was an inversion episode which started in the late cretaceous. after that inversion the kattegat became an area of non-deposition and net-erosion until net-sedimentation was resumed in the saalian. at that time the basin floor of the kattegat was characterised by strongly undulating relief controlled by largeand small-scale structures in the pre-quaternary basement. the survey area can be divided into several subareas, based on bathymetry and seabed sediment types (fig. 2). the northern part of the survey area is smooth, with sand and silt. the central and southern parts show ridges with a relief of more than 1 m. it is suggested that the ridges were formed by waves at a time of lower than present sea level (leth et al. 2009). the area with the highest density of gravel and boulders corresponds to the shallowest part of the central survey area. the boulders are generally located in arc-like, nw– se-striking, narrow structures found mainly in the southern and western parts of the survey area (figs 1, 2). depths over 18 m characterised by sand and silt with pebbles occur near the western and eastern margins of the survey area. pre-quaternary strata – unit pq the pre-quaternary surface is a regional erosional unconformity with high amplitude seismic reflection. the reflector has been mapped throughout the area down to the limit of the penetration of the sparker system at c. 60 m below sea level (b.s.l.; figs 3, 4d). the pre-quaternary surface lies deeper than 60 m b.s.l. in the eastern part of the area. two boreholes penetrated several metres of pre-quaternary silty, fine sand. analysis of palynomorphs in two samples yielded an upper cretaceous age (k. dybkær and e. sheldon, personal communication 2009). glacial deposits – unit gl the glacial deposits have been divided into two subunits, gl1 and gl2. the lower subunit gl1 is found in the central and southern parts of the survey area. it shows a characteristic mediumto low-amplitude, parallel, wavy and chaotic seismic pattern. the parallel or wavy pattern is interpreted as representing undisturbed sorted and layered sediments, whereas the wavy or chaotic pattern is interpreted as representing glacially dislocated sediments. a unit with a similar seismic character as gl1 has been found in various parts of the kattegat, and has been referred to the late saalian, eemian and middle weichselian (vangkilde-pedersen et al. 1993). marine and glaciogene sediments of these ages have been recorded in sediment cores from other parts of the kattegat region (larsen et al. 2009). the transition to seismic subunit gl2 is sharp in the southern part of the survey area, whereas it is more gradual in the central part. internally gl2 shows a mediumto highamplitude, seismic facies pattern that is chaotic, mounded and channelled. low-angle oblique reflectors cutting through the whole subunit are interpreted as large-scale, glaciotectonic deformation structures. data from sediment samples show that the subunit mainly consists of sand with poorly and wellsorted layers of clay, silt and gravel. a high density of cobbles and boulders is seen where gl2 crops out on the sea floor, as confirmed by surface samples (fig. 4a). a unit showing a similar seismic pattern and with similar deposits from the lysegrund area (fig. 1) has been interpreted as subglacial and glaciofluvial deposits. the gl2 subunit probably corresponds to the m3 unit at lysegrund, which consists of icemargin sediments deposited during the retreat stage of the main advance (novak 1996). the seismic pattern, facies association, unit morphology, sea-floor character and lithology all suggest that gl2 represents similar ice marginal deposits. 2 km gl2 lg1 lg2 pq gl1 fault h2 gl2 gl1 lg1 h1 gas no data north south lg2 bh-05bh-01bh-02bh-09 40 50 60 d ep th ( m b el o w s ea l ev el ) 30 20 pq fig. 3. schematic model of the seismic units based on a n–s-trending section along utm 634000 me (wgs84). the unit names refer to descriptions in the text. the boreholes (bh-01, bh-02, bh-05 and bh-09) are located at distances from 400 m to 1.5 km from the profile (fig. 2). 33 late glacial deposits – unit lg the distribution of the late glacial seismic unit lg is governed by the morphology of the underlying glacial surface (fig. 4c). its maximum thickness is 45 m. towards the south the depressions in the glacial surface are characterised by shallow channels and small basins, and it is possible to correlate these depressions to a system of elongated fault-related basins in the pre-quaternary surface sw of anholt (binzer & stockmarr 1994). the seismic unit lg is subdivided into two subunits, lg1 and lg2, with a gradational boundary. lg1 shows an external apron or mound morphology with internal composite mounds as well as a hummocky, shingled, parallel reflection pattern. at its base, lg2 shows an onlap-downlap, draping style. upwards it shows a gradually decreasing amplitude and a more pronounced semi-transparent, parallel seismic facies. in general, lg2 terminates upwards into an erosional unconformity. a seismic ‘blacking out’ area internally in lg2 indicates gas content in a discrete level associated with a pronounced reflector that probably represents a sealing clay layer. data from boreholes show that unit lg consists of a fining-upward sequence with sand and gravel at its base and layers of clay with sand and silt laminae towards its top. the characteristic seismic expressions of unit lg have also been recorded from other parts of southern kattegat. for instance, jensen et al. (2002) reported two stages in a late glacial unit located in elongated fault-related basins and suggested that the two stages are related to re-activation of normal fault activity in the elongated depressions in the period from 15 to 13.5 calendar ka bp. holocene deposits – unit h the transition from the late glacial to the holocene unit h is seen as a shift to high-amplitude reflectors. truncation of the rhythmic parallel facies of lg2 is succeeded by mounds, hummocky oblique and sub-parallel reflector patterns in the holocene subunit h1. this subunit was previously referred to the late glacial (leth et al. 2009), but after reassessment of the abrupt changes in the seismic signature and its distribution we conclude that the subunit is of early holocene age. in the southern part of the survey area, 400–600 m wide channels orientated wsw–ene and ssw–nne are filled with late glacial and early holocene (h1) deposits. the base level of h1 in these channels is around 22 m b.s.l. h1 is deposited above the truncated lg2 unit and its distribution is confined by the older lg basins. the same base level is found in wider areas in the north and is likewise unconformable to the underlying lg unit. a significantly deeper channel crosses the central survey area with internal seismic structures that indicate a unidirectional flow from west to east (h2; fig 3). the base of this channel is generally at 26–28 m b.s.l. locally the channel widens to 1000 m. organic-rich sediments of holocene age representing a lowstand at 35 m b.s.l. are well-known from the kattegat area. after a fluvial event west of lysegrund the water level stabilised at 34 m b.s.l. (novak & björk 1998). lagoonal deposits overlying truncated, rhythmic, late glacial clay-sand layers are found south-east of the survey area (bennike et al. 2000; novak & pedersen 2000). in a major area these sediments are mostly found between 35 and 24 m b.s.l. and have been dated to the early holocene (bennike et al. 2000). da cb 25–30 m 30–35 m 35–40 m 40–45 m 45–50 m 50–55 m 55–60 m no data 5 km 56°33´n 11°15´e gravel-pebbles 25–100% boulders sand-pebbles 1–25% boulders sand and silt, pebbly solitary boulders sand and silt 26–28 m 24–26 m 22–24 m 20–22 m 18–20 m <18 m 20–25 m 25–30 m 35–40 m >45 m 30–35 m 40–45 m fig. 4. maps showing seabed sediments and three seismic stratigraphic levels. a: seabed sediment types. b: depth of base holocene. black line indicates fault. c: depth to the top of the glacial deposits. the dashed line shows the disturbed–undisturbed gl1 interface (see text). d: pre-quaternary surface morphology. black line indicates fault. 3434 lagoonal sediments at 18 m b.s.l., found 30 km east of the study area, were also dated to the early holocene (novak & pedersen 2000). at most sites h2 only represents a veneer of fineto coarse-grained sand with gravel, occasionally with silt and clay laminae as well as shell fragments of marine molluscs. however, in one of the samples marine shells and organic material are found at 9 m below the sea floor. it is suggested that the early holocene transgression reached a level high enough to submerge the survey area prior to 9.9 calendar ka bp. a channel in the central part of the survey area drained towards the east. structural features in the central survey area, reflectors in the sparker profiles indicate listric normal faults dipping south-wards (fig. 5). a 2.9 km long, e–w-striking sea-floor lineament (fig. 4b, c) represents the top of the headwall scar, and the fault has a significant signature through the whole quaternary package. the location and strike of the faults follow the structures in the pre-quaternary basement (binzer & stockmarr 1994), which indicates that they represent re-activations of old faults. final remarks based on geophysical data, boreholes and cone penetration tests we have documented that the geological architecture of the study area is very complex. the pre-quaternary (late cretaceous) basement is overlain by two glacial, two late glacial and two holocene subunits. the widespread late glacial deposits that are up to 45 m thick are dominated by fine-grained sediments, in some areas with gas. the distribution of the holocene deposits indicates the presence of channels that drained into the deeper part of the kattegat during the early holocene. the data described here are of great importance to the geotechnical evaluation prior to the planned foundation of windmills. acknowledgements we thank energinet.dk for permission to use the above data and to publish the geological results from the anholt wind-farm project. references bennike, o., jensen, j.b., konradi, p.b., lemke, w. & heinemeier 2000: early holocene drowned lagoonal deposits from the kattegat, southern scandinavia. boreas 29, 272–286. binzer, k. & stockmarr, j. 1994: geological map of denmark, 1:500 000. pre-quaternary surface topography of denmark. danmarks geologiske undersøgelse kortserie 44, 10 pp., 2 maps. jensen, j.b., petersen, k.s., konradi, p., kuijpers, a., bennike, o., lemke, w. & endler, r. 2002: neotectonics, sea-level changes and biological evolution in the fennoscandian border zone of the southern kattegat sea. boreas 31, 133–150. larsen, n.k., knudsen, k.l., krohn, c.f., kronborg, c., murray, a.s. & nielsen, o.b. 2009: late quaternary ice sheet, lake and sea history of southwest scandinavia – a synthesis. boreas 38, 732–761. leth, j.o., alhamdani, z., novak, b., barzani, s.m. & hindrichsen, c. 2009: anholt offshore wind farm. marine geophysical investigations. danmarks og grønlands geologiske undersøgelse rapport 2009/45, 411 pp. lykke-andersen, h., knudsen, k.l. & christiansen, c. 1993: the quaternary of the kattegat area, scandinavia: a review. boreas 22, 269–281. novak, b. 1996: en maringeologisk undersøgelse af kvartære lag på lysegrund, sydlige kattegat, danmark. geologisk tidsskrift 2, 21–25. novak, b & björck, s. 1998: marine seismic studies in southern kattegat, with special emphasis on longitudinal bars and their possible relationship to the drainage of the ancylus lake. gff 120, 297–306. stockholm: geological society of sweden. novak, b. & pedersen, g.k. 2000: sedimentology, seismic facies and stratigraphy of a holocene spit-platform complex interpreted from high-resolution shallow seismics, lysegrund, southern kattegat, denmark. marine geology 162, 317–335. vangkilde-pedersen, t., lykke-andersen, h. & lind, g. 1993: dislocated quaternary deposits in southeastern kattegat – a glacial or gravitational phenomenon? boreas 22, 329–336. 20 40 60 tw o -w ay t ra ve l ti m e (m se c) nn ssn s 100 m multiplemultiplemultiple 15 30 45 d ep th (m ) fig. 5. part of the n –s-orientated seismic profile aw039 showing a fault that dips to the south. a ridge is seen along the top of the fault. note the subsidence of the seabed above the fault. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: jol@geus.dk geological survey of denmark and greenland. bulletin 10, 37-40 visual inspection and comparison of lithological, topographical, geochemical and geophysical maps is a necessary (and often successful) way of interpreting geological features across poorly exposed or poorly explored areas. in mineral exploration, geochemical and geophysical maps are used to visually identify anomalies believed to reflect mineral occurrences. outcropping mineral occurrences of a certain size can produce elevated concentrations of elements in stream sediment samples, or create magnetic or radioactive signals retrievable by airborne surveys. however, if the regional data are widely spaced, and if the occurrences are small or poorly exposed, the anomalies created may be too subtle to be recognised visually. in such cases, statistical data analyses may help identify deviations from background variations and trends in the data. furthermore, visually based correlation between distribution patterns on maps is often limited to a few parameters, and multi-variable relationships are easily overlooked or not possible to deduce. gold occurrences in greenland occur in specific host rocks as a result of mineralising processes, e.g. hydrothermal activity, that have affected a larger area. thus, the favourable environment for a gold occurrence may be reflected in regional scale survey data as subtle changes in certain geochemical and geophysical properties. if such changes, i.e. a multi-parameter signature for gold mineralisation, can be identified, the regional data can be searched for areas that might hold the same signature. this paper presents results gained in the application of a statistical spatial analysis method to investigate how known gold showings are expressed in a range of regional datasets from the nuuk region, southern west greenland. the main objective is to identify new areas with gold potential, and to an integrative and quantitative assessment of the gold potential of the nuuk region,west greenland bo møller stensgaard,thorkild m. rasmussen and agnete steenfelt © geus, 2006. geological survey of denmark and greenland bulletin 10, 37–40. available at: www.geus.dk/publications/bull 37 64°30'n 65°n 52°w 51°w isukasia fis ke fjo rd qq aa q a 50°w se rm its ia q inland ice davis strait go dt hå bs fjo rd godthåbsfjord nuuk ameralik qooqqut storø bjørneøen ivisaartoq greenland 25 km gneiss ice amphibolite metasediment ultramafic rock fault and shear zones late archaean granites isua gold group most favourable areas for: bjørneøen gold group storø gold group storø showings bjørneøen showings isua showings unpredicted showings insignificant showings to form a group fig. 1. simplified geological map with 137 km2 (1% of the analysed area) outlined as most favourable for each of the three groups of gold showings (isua, bjørneøen and storø groups). outlines of the rock units are from digital versions of 1:100 000 and 1:2 500 000 scale geological maps published by the survey. q, qingaaq; a, aappalaartoq. quantify the signatures of the showings. the first tests based on this approach were reported in nielsen et al. (2004) and steensgaard et al. (2006). compilation of available information in databases and company reports the nuuk region is underlain by early to late archaean rocks dominated by quartzofeldspathic orthogneisses, but with a significant proportion of supracrustal belts, some of which contain greenstones (fig. 1). the belts represent remnants of archaean volcanic and associated intrusive rocks, with subordinate chemical and clastic sediments. recent studies have revealed that the belts represent a number of individual successions formed at different locations and ages. the most common rock type is amphibolite of tholeiitic to komatiitic composition, but andesitic compositions have also been recognised in some belts (hollis et al. 2005, 2006, this volume). metasedimentary rocks are mostly garnet-biotite schist. the supracrustal rocks host several gold occurrences in association with hydrothermal alteration (fig. 1; appel et al. 2005). two of these, at qingaaq and aappalaartoq on storø (q and a, fig. 1) are currently targets for commercial drilling by nunaminerals a/s. datasets analysed twenty-nine datasets have so far been included in the statistical analysis. the data are derived from regional scale surveys, and comprise stream sediment geochemistry and aeromagnetic data. the geochemical data include concentrations of al2o3, as, au, cao, cs, fe2o3, k2o, la, mgo, na2o, p2o5, rb, sb, sio2, th, tio2, v, zr, zn, u, ni/mgo and ni + cr in the < 0.1 mm grain size fraction of stream sediments. the magnetic data include the total magnetic field intensity (tmi), and derivatives of the total field, i.e. vertical gradient of tmi, horizontal gradients of tmi in various directions (hgtmi), and the amplitude of the hgtmi. a search through available non-confidential data reported by exploration companies, as well as survey databases, provided 38 data n o . o f sh o w in gs showings s s s no showings empirical distribution functions are calculated for different datasets – one for 'showings', one for 'no showings' ? ?? ? ? ? ? ? ? ? ? ? ? ? ?? ? ? ? ? ? ? ? ? ? ? ?? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ?? ? ? ? ? ? ? ? ? ? ? ?? ? ? ? ? ? ? ? ? ? ? ? ? ? quantified signatures in different datasets a search for similar signatures is carried out. the degree of similarities is used to construct mineral potential maps mineral potential map highlow favourability for showings geoscientific data sets are pixelated 2 1 4 3 fig. 2. illustration of the statistical approach. 1: regional datasets and gold showings (s) are referred to a common grid with fixed cell size. 2: empirical distribution functions for cells with and without known gold showings are used to establish the signatures of the showings in each dataset. 3: likelihood ratio functions for all datasets are calculated as the ratio of the two previous functions, expressing the likeliness of a showing to be present,. groupings of gold showings are based on constructed mineral potential maps for each showing. pixel by pixel, these maps outline the predicted favourability for data signatures similar to the selected showing, i.e., the potential for a similar showing to be present. in this way the mutual prediction capability amongst different showings can be compared. 4: finally, mineral potential maps are calculated for each identified group of showings. these maps are based on the signatures found to be indicative for each of the different groups. all known localities from which gold-bearing rock samples had been collected. methodology the principles of the methodology are illustrated in fig. 2. all spatial datasets must be converted into a form where they can be compared and subjected to statistical analysis. each dataset is presented as a regular grid, using a common cell size of 200 m x 200 m, and a gridding procedure based on principles of minimum curvature. note that the resulting grid values for geochemical and geophysical data are obtained by interpolation, as most measured points of the original datasets have a larger spacing than 200 m. there is therefore a risk that interpolated values in some areas misrepresent the data. grid cells with and without gold showings are registered. in this study, a gold showing is defined as a cell in which rock samples with 1 ppm au have been collected. in total, 52 gold showings were identified. the data were analysed using statistical methodologies based on chung (2003, and references therein). in summary, analysing the integrated data signature of cells with gold showings, and using this information to calculate for each cell the probability that it contains a similar gold showing, can be used to construct mineral potential maps. a mineral potential map defines zones or classes of different favourability for a certain mineralisation type. the methodologies involve calculation of empirical distribution functions (edfs) for cells with and without gold showings, which can be regarded as the signature of the gold showings and the regional background (fig. 3). the ratio between the two edfs is used to calculate likelihood ratio functions (lfs) for each dataset. a joint likelihood ratio function (jlf) is then calculated as the product of the lfs. this jlf is then used to construct the mineral potential map. grouping of gold showings in our first simple approach, we treated all known gold showings as one group (nielsen et al. 2004). however, the results suggested that some gold showings had different signatures, and statistical methods were subsequently used to test whether the showings could be grouped according to their signatures. a gold potential map was then made for each single showing, based on the datasets that are characteristic of all showings. then, the capability of the selected showing to predict the remaining showings was evaluated. as a result, a group defined this way consists of showings that have mutual prediction capabilities. the analysis resulted in the identification of three main groups of gold showings (fig. 1): the storø group (11 showings in a discrete area at qingaaq and aappalaartoq), the bjørneøen group (10 showings located around central bjørneøen), and the isua group (19 widely spread showings at isukasia and one showing at ivisaartoq). two other groups shown on fig. 1 are not treated further here, namely a group of eight showings without mutual prediction capabilities with other showings, and another group of four showings which is considered too small to be statistically significant. gold potential maps figure 1 displays a gold potential map, where red, blue and green boundaries outline the most favourable areas for finding gold occurrences according to the criteria established for each of the three main groups. it is not surprising that areas immediately surrounding known gold showings are predicted as favourable, but more interestingly, a number of favourable areas are also predicted outside known gold showings. many of the latter areas are within or adjacent to supracrustal rocks, that may otherwise be regarded as potential hosts to gold. some of the most notable favourable areas are east of the fjord qooqqut, areas on storø favourable for bjørneøen group gold showings, and areas on south-eastern sermitsiaq and farther to the southwest. finally, several favourable areas are outlined along fiskefjord (fig. 1). this area has hardly been prospected for gold, but comprises both supracrustal rocks and structures that are similar to the hosts of some of the known gold showings. 39 fig. 3. single-parameter signatures of three groups of gold showings and the background. a: arsenic (as), based on concentrations in the fine fraction of stream sediment. b: nickel/magnesium (ni/mg) ratio, based on concentrations in the fine fraction of stream sediment. c: caesium (cs), based on concentrations in the fine fraction of stream sediment. d: vertical gradient (vgtmi), based on airborne measurements of total magnetic field intensity. vertical gradient of total magnetic intensity field [nt/m] -0.6 -0.4 -0.2 0.0 0.2 0.4 0.6 n o rm al is e d fr e q u e n cy 0.0 2.0 4.0 6.0 8.0 d cs [ppm] 6 n o rm al is ed fr eq u en cy 0.0 0.8 1.6 2.4 3.2 3 4 52 c ni/mg ratio 5 15 25 35 45 55 n o rm al is ed fr eq u en cy 0.00 0.04 0.08 0.12 0.16 b as [ppm] 2 6 10 14 18 22 26 30 n o rm al is ed fr eq u en cy 0.0 0.2 0.4 0.6 0.8 a isua group bjørneøen group storø group background 40 when a mineral potential map for gold is evaluated, it is important to keep in mind that areas outlined as having a low favourability need not necessarily be discarded as targets for gold since the constructed maps only reflect the data and known gold occurrences included in the analysis. signatures of gold showings the systematic construction of signatures of showings and background (edfs) draw attention to the significance of parameters that are not immediately or traditionally regarded as indicative of gold mineralisation. this is illustrated in fig. 3 by four examples of single-parameter empirical distribution functions for each of the three groups of gold showings, together with the background variation. it has previously been established that the combination of elevated as in stream sediment and gold mineralisation has not been found outside storø (fig. 3a). however, the possible significance of the ni/mg ratio (fig. 3b) in relation to gold has not previously been substantiated in a quantitative way. this observation is an incentive to conduct further studies of host rock properties in terms of their ni and mg behaviour. the rare and unique as signature of the storø group poses a problem in the search for areas holding a potential for new showings. since this signature only exists on storø, only this area will be outlined as favourable for the storø group. consequently, the as data are omitted in the calculation of cells with a potential for storø-type showings. steenfelt (2000) has argued that cs may be considered a pathfinder element for gold mineralisation associated with granite-related hydrothermal alteration. the signature for cs in stream sediment confirms that the gold at storø and bjørneøen is indeed located in a cs-rich environment (fig. 3c), but also suggests that the mineralisation at isua has a different nature. the last example (fig. 3d) illustrates the indicated significance of the vertical gradient of the total magnetic field intensity (denoted vgtmi). this signature for the storø and bjørneøen groups (–0.2 to 0.1 nt/m) is identical to the background signature, and is consequently regarded as non-indicative for these groups. the isua group has a more dispersed signature (–0.5 to 0.8 nt/m), with a more highly indicative positive and negative vgtmi than the other groups. this probably reflects local scale changes in lithology. other datasets, which yield a characteristic signature of one or more of the three groups of gold showings, are the au, cs, rb, la, th, and u concentrations in stream sediment geochemistry. further work and perspectives the results obtained in the first tests of the applied statistical method have been positive in the sense that areas with high probability for new gold occurrences were outlined in geological settings that would otherwise be regarded as prospective for gold. in the next phase of the study additional datasets will be included, such as airborne gamma-ray spectrometry, digital topography, lithology and structure; the gridding procedure will be refined, and the influence of changing the cell size and the use of other statistical methods will be tested. this kind of statistical approach has a potential for a range of applications, such as the prediction of other types of mineral occurrences, and identification and mapping of specific rock units in poorly exposed or poorly known terrain. multivariate statistical analysis seems inevitable as a way to optimise the use of geoscientific data. however, it is important that geoscientists critically evaluate the results of any statistical analysis in order to reveal misleading effects of interpolation and other ways of standardising parameters that by their nature are very variable. acknowledgement the bureau of minerals and petroleum, government of greenland, financially supported the project. references appel, p.w.u., coller, d., coller, v., heijlen, w., moberg, e., polat, a., raith, j., schjøth, f., stendal, h. & thomassen, b. 2005: is there a gold province in the nuuk region? danmarks og grønlands geologiske undersøgelse rapport 2005/27, 79 pp., 1 cd-rom. chung, c.f. 2003: use of airborne geophysical surveys for constructing mineral potential maps. economic geology monograph 11, 879–891. hollis, j.a., van gool, j.a.m., steenfelt, a. & garde, a.a. 2005: greenstone belts in the central godthåbsfjord region, southern west greenland. geological survey of denmark and greenland bulletin 7, 65–68. hollis, j.a., frei, d., van gool, j.a.m., garde, a.a. & persson, m. 2006: using zircon geochronology to resolve the archaean geology of southern west greenland. geological survey of denmark and greenland bulletin 10, 25–28. nielsen, b.m., rasmussen, t.m. & steenfelt, a. 2004: gold potential of the nuuk region based on multi-parameter spatial modelling of known gold showings. danmarks og grønlands geologiske undersøgelse rapport 2004/121, 155 pp. steenfelt, a. 2001: geochemical atlas of greenland – west and south greenland. danmarks og grønlands geologiske undersøgelse rapport 2001/46, 39 pp. steensgaard, b.m., steenfelt, a. & rasmussen, t.m. 2006: gold potential of the nuuk region based on multi-parameter spatial modelling. progress 2005. danmarks og grønlands geologiske undersøgelse rapport 2006/27, 207 pp. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: bmst@geus.dk geological survey of denmark and greenland bulletin 6, 77-93 77geological survey of denmark and greenland bulletin 6, 77–93 © geus, 2004 the eleonore sø and målebjerg foreland windows, east greenland caledonides, and the demise of the ‘stockwerke’ concept a.k. higgins and a. graham leslie recognition of the eleonore sø and målebjerg foreland windows during the 1997–1998 regional mapping expeditions to the east greenland caledonides provided critical evidence for largescale, westward-directed thrusting in the kong oscar fjord region (72°–75°n), a revelation that dealt a final blow to the ‘stockwerke’ concept of an in situ highly mobile infrastructure characterised by rising fronts of caledonian migmatisation and metasomatism. this paper reviews earlier investigations in both the eleonore sø and målebjerg areas, and the misinterpretations of rock units that initially obscured recognition of their foreland affinity. the eleonore sø and målebjerg windows can now be placed in context, as part of the lowest structural level of the foreland-propagating thrust pile of the kong oscar fjord region. keywords: caledonides, east greenland, tectonic windows a.k.h., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. email: akh@geus.dk a.g.l., british geological survey, murchison house, edinburgh eh9 3la, uk. the 1300 km long east greenland caledonides (70°– 81°30′n) can be broadly divided into western marginal and eastern thick-skinned thrust belts (fig. 1; higgins & leslie 2000; higgins et al. 2001, 2004). the western thrust margin of the orogen against the caledonian foreland is largely obscured by the greenland inland ice, with the most continuous foreland exposures west of kronprins christian land in the extreme north. elsewhere foreland areas are locally preserved in the westernmost nunataks and in scattered tectonic windows exposed along the length of the marginal thrust belt (fig. 1). while the foreland windows all exhibit some disturbance due to caledonian deformation, and have therefore been classified as parautochthonous, the similarities of the successions preserved within the various windows and that in the undisturbed foreland, suggest they are only slightly displaced from their original locations (higgins et al. 2001). prior to the survey’s 1997–1998 regional mapping programme, large-scale thrusting had not been demonstrated in the kong oscar fjord region (fig. 2; 72°–75°n). indeed, new investigations in this region (hartz & andresen 1995; andresen & hartz 1998; andresen et al. 1998), had led to interpretations of caledonian orogenesis in terms of upward and lateral movement of light, low viscosity, lower crustal material towards the region of maximum crustal extension, a process compared to haller’s (1953, 1970, 1971) ‘stockwerke’ concept, and that carried the implication that orogenic contraction was negligible. the recognition of the målebjerg and eleonore sø foreland windows during the survey’s 1997–1998 regional mapping provided incontrovertible evidence for large-scale westward-directed caledonian thrusting in the kong oscar fjord region (figs 2, 3). this discovery completely undermined arguments for explaining orogenic development in terms of the in situ ‘stockwerke’ concept or similar processes, and at the same time resolved a number of outstanding problems of east greenland geology. the areas of both geus bulletin 6.pmd 10-02-2005, 09:5477 78 ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ c a le d o n id e s greenland 70°n 74°n 78°n 82°w 35°w wandel sea centrumsø 100 km jameson land scoresby sund 25°w traill ø wollaston forland bessel fjord danmarkshavn lambert land peary land station nord kr on pr ins c hr ist ian l an d nørreland window kronprins christian land thin-skinned thrust belt (parautochthonous foreland) palaeogene basalts palaeogene intrusions wandel sea basin: carboniferous–palaeogene sediments east greenland basins: carboniferous–cretaceous sediments devonian – continental sediments late to post-kinematic granites neoproterozoic–ordovician sediments (east greenland) neoproterozoic–silurian sediments (eastern north greenland) palaeo-mesoproterozoic sediments and basalts (eastern north greenland) crystalline complexes and sediments (archaean–mesoproterozoic) neoproterozoic–silurian sediments (north greenland) palaeo-mesoproterozoic sediments and basalts (eastern north greenland) mainly crystalline rocks – parautochthonous windows thrust fault/shear zone tectonic zone boundary post-caledonian late to post-caledonian caledonian orogenic belt caledonian foreland ▲ ▲ fig. 2 målebjerg window gåseland window c a le d o n ia n so le t h ru st in la n d ic e t h r u s t b e lt t h ic k s k in n e d t h r u s t b e lt charcot land window eleonore sø window hamberg gletscher foreland dronning louise land ▲ ▲ ▲ m a r g in a l fig. 1. geological map of the east greenland caledonides, showing location of the foreland windows in the western marginal thrust belt. the frame indicates the region between 71°50′ and 74°30′n, shown at a larger scale in fig. 2, which includes the målebjerg and eleonore sø windows. modified from higgins & leslie (2000). geus bulletin 6.pmd 10-02-2005, 09:5478 79 ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i ▲ ▲ ▲ ▲ ▲ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ fjd eleonore sø window hamberg gletscher foreland målebjerg window charcot land window nathorst land hudson land ella ø suess land lbl pd p andrée land lyell land fr f stauning alper kong oscar fjord kfjf es ml m cecilia nunatak fig 3 fjd neoproterozoic–ordovician post-caledonian caledonian granite 930 ma granite palaeoproterozoic–ordovician krummedal sequence krummedal sequence palaeoproterozoic orthogneiss archaean–palaeoproterozoic orthogneiss ▲ ▲ ▲ ▲ franz joseph allochthon foreland hagar bjerg thrust sheet niggli spids thrust sheet detachment extensional fault thrust cross-section line ▲ ▲ i i ■ ■ ■ ■ 72°72° 73° 74° 27° 100 km fig. 2. geological map of north-east greenland 71°50′– 74°30′n, showing location of the eleonore sø, målebjerg and charcot land windows, and the hamberg gletscher foreland. the legend depicts the units contained in the two thrust sheets and franz joseph allochthon overlying the windows. the målebjerg and eleonore sø areas are shown in more detail in figs 5 and 6. es, eleonore sø; fjd, franz joseph detachment; frf, fjord region fault; kfjf, kejser franz joseph fjord; lbl, louise boyd land; m, målebjerg; ml, j.l. mowinckel land; p, petermann bjerg; pd, petermann detachment. the line of the cross-section shown in fig. 3 is indicated. geus bulletin 6.pmd 10-02-2005, 09:5479 80 windows had been investigated prior to the survey’s regional mapping, the eleonore sø area by katz (1952) and the målebjerg area in andrée land by haller (1953). however, while haller (1971) had speculated that the eleonore sø region might be a window (see below), neither katz nor haller identified the lower palaeozoic rock units whose presence clarifies the structural setting beyond any doubt. this paper reviews the history of investigations in the eleonore sø and målebjerg areas, and the early misinterpretations of rock units that obscured their recognition as parts of the caledonian foreland. these misinterpretations were closely linked to the evolution of ideas to explain orogenic developments in east greenland, that culminated in the ‘stockwerke’ concept as elaborated by haller (1970, 1971). haller’s ‘stockwerke’ model was a development of the earlier ideas of backlund (1930, 1933) and wegmann (1935). the crystalline gneiss complexes constituting the central metamorphic complex that underlie, and appear to be interleaved with high grade metasediments, were envisaged by haller as elements of a highly mobile infrastructure formed by the rise of caledonian fronts of migmatisation and metasomatism. associated mechanical and chemical changes were thought to have led to in situ transformation of a succession of sedimentary rocks into the gneissic and granitic rocks of the infrastructure, which was considered ‘entirely rejuvenated’. the mobile migmatite domes of the infrastructure were bordered by and overlain by the more rigid sedimentary suprastructure, with the two levels separated by a thick, often strongly folded, ‘zone of detachment’. the term ‘stockwerke’ refers to the different levels of the growing orogenic belt; in german ‘stockwerke’ refers to the floors or stories of a house. in haller’s map compilations the infrastructure of the central metamorphic complex is depicted as caledonian synorogenic granite (koch & haller 1971), and he remarks that regional thrusting was probably of “no importance in their formation” (haller 1971, p. 179). the spectacular ‘stockwerke’ structures in east greenland, and haller’s drawings, are still presented in modern textbooks (e.g. best 2003) as classic examples of mantled gneiss domes, and metamorphic core complexes. the målebjerg and eleonore sø windows can now be placed in their correct context as parts of the lowest structural level of the foreland-propagating thrust pile in the southern half of the caledonian orogen (70°–75°n; elvevold et al. 2000; higgins et al. 2004). these foreland windows are structurally overlain by a lower niggli spids thrust sheet and an upper hagar bjerg thrust sheet, both with substantial westward displacements (fig. 3). the very thick neoproterozoic to lower palaeozoic succession (eleonore bay supergroup, tillite group, kong oscar fjord group) is distinguished as the franz joseph allochthon, and viewed as an upper detached part of the hagar bjerg thrust sheet. hamberg gletscher foreland boyd bastion fault eleonore sø window målebjerg window fjord region fault 10 km ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ■ ■ ■ ■ ■ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ■ ■ caledonian granite 930 ma granite neoproterozoic–ordovician krummedal sequence krummedal sequence palaeoproterozoic orthogneiss archaean–palaeoproterozoic neoproterozoic–ordovician palaeoproterozoic ▲ ▲ ▲ ▲ ▲ ▲ ■ ■ extensional fault detachment thrust ■ ■ franz joseph allochthon   foreland       hagar bjerg thrust sheet    niggli spids thrust sheet fig. 3. cross-section through the eleonore sø and målebjerg windows showing the foreland windows overlain by two thrust sheets and the franz joseph allochthon. section line is indicated on fig. 2. geus bulletin 6.pmd 10-02-2005, 09:5480 81 geological setting the eleonore sø and målebjerg windows (figs 2, 3) are characterised by thin (< 400 m) neoproterozoic – lower palaeozoic sedimentary successions. the latter comprise a lens-like 31 m thick diamictite in the målebjerg window correlated with the vendian tillite group, a 143–350 m thick quartzite sequence with skolithos ichnofossils of early cambrian age defined as the slottet formation, and a 32–45 m thick dolomite sequence of cambrian–ordovician age defined as the målebjerg formation (smith et al. 2004, this volume). this thin vendian–ordovician foreland succession is in great contrast to the 18.5 km thick, partly equivalent succession, preserved in the structurally overlying franz joseph allochthon. in the eleonore sø window the slottet formation overlies a palaeoproterozoic sedimentary-volcanic assemblage with profound unconformity, and in the målebjerg window unconformably overlies gneisses of presumed palaeoproterozoic age. the niggli spids thrust sheet structurally overlying both windows (fig. 3) incorporates crystalline gneiss complexes and high-grade metasedimentary successions, which were reworked to varying degrees during caledonian orogenesis. ion microprobe studies on zircons from the orthogneisses have yielded archaean and palaeoproterozoic protolith ages (thrane 2002; unpublished data 2004, f. kalsbeek and a.p. nutman), that confirm earlier less precise isotopic ages by other methods (e.g. rex & gledhill 1981). the highgrade metasedimentary rocks of both the niggli spids thrust sheet and the higher hagar bjerg thrust sheets are correlated with the krummedal supracrustal sequence (higgins 1988). the krummedal sequence metasedimentary rocks in the hagar bjerg thrust sheet host a suite of 940–910 ma augen granites generated during an early neoproterozoic thermal event (kalsbeek et al. 2000; watt et al. 2000; watt & thrane 2001). ion microprobe studies of detrital zircons from the krummedal sequence show the youngest detrital zircons are about 1050 ma old, and deposition of the sediments must therefore have taken place in the period c. 1050–940 ma ago (late mesoproterozoic – early neoproterozoic). high-grade regional metamorphism and associated anatexis during the caledonian orogeny led to generation of a new suite of 440–425 ma granites (watt et al. 2000; hartz et al. 2001; kalsbeek et al. 2001a, b). the franz joseph allochthon is made up of the c. 13 km thick neoproterozoic eleonore bay supergroup (riphean–sturtian; post-900 to c. 590 ma), the 800– 1000 m thick vendian tillite group (hambrey & spencer 1987), and the c. 4 km thick cambrian–ordovician kong oscar fjord group (cowie & adams 1957; smith & bjerreskov 1994; smith et al. 2004, this volume). this succession is widely exposed in the central fjord system of the kong oscar fjord region, and also occurs in a more restricted area to the west in louise boyd land and around petermann bjerg (fig. 2). the contact between the eleonore bay supergroup and underlying high-grade krummedal metasediments of the hagar bjerg thrust sheet is a shear zone, in which both extensional and contractional strain have been recorded. in the west the shear zone is known as the petermann detachment (pd, fig. 2; escher & jones 1998, 1999), and in the east the franz joseph detachment (fjd, fig. 2). the latter is only well exposed between northern andrée land and hudson land (leslie & higgins 1998, 1999), and between lyell land and the stauning alper (cf. tindern detachment of white et al. 2002). in most of the fjord region the present-day west limit of the eleonore bay supergroup outcrop is a late-orogenic extensional fault, the ‘fjord region fault’ of this paper (frf, fig. 2). this corresponds to the ‘fjord zone fault’ of larsen & bengaard (1991) and to part of the ‘fjord region detachment’ system of hartz & andresen (1995) and co-workers (andresen et al. 1998; hartz et al. 2000; white et al. 2002). conodonts extracted from the uppermost levels of the ordovician succession exhibit very low conodont alteration indices (smith 1991; stouge et al. 2002), which demonstrate that the franz joseph allochthon in this region cannot have been over-ridden by higher thrust sheets. historical review: the målebjerg and eleonore sø areas the investigations of john haller and others working with lauge koch’s long series of east greenland geological expeditions (1926–1958) more than 40 years ago, essentially predated the era of isotopic age determinations. the first, very few, k-ar ages from east greenland only became available in 1961 (haller & kulp 1962), after the cessation of field work in 1958. the evolution of ideas and the conclusions of their studies were thus almost entirely based on field observations and interpretations. the revolutionary concept of plate tectonics did not make its mark until the late 1960s, and while it was widely accepted by many geus bulletin 6.pmd 10-02-2005, 09:5481 82 geologists, john haller considered that its enthusiastic reception had obscured its shortcomings (see haller 1979). lauge koch’s sledge journeys along the length of the caledonian orogen led him to propose that the greater part of the gneisses that previous expeditions assumed to be archaean were “in reality the nucleus of a caledonian folding range” (koch 1929, p. 60, fig. 20). however, at this time the gneisses of the inner fjords between scoresby sund and kejser franz joseph fjord were still considered to be archaean. meanwhile, british geologists working in the inner part of the kong oscar fjord region compared the gneisses and metasedimentary rocks of the so called ‘central metamorphic complex’ to the lewisian and dalradian of scotland, and concluded that the caledonian orogeny was ‘superficial’ (wordie 1930; parkinson & whittard 1931). helge backlund investigated the same region, but reached a different interpretation (backlund 1930, 1932). he proposed that wordie’s ‘archaean’ granites and gneisses were the result of caledonian granitisation and migmatisation of a varied sedimentary succession, similar to the processes of gneiss formation in fennoscandia. backlund’s views were largely supported by wegmann’s (1935) report on the ‘caledonian orogeny’, in which wegmann speculated that any former basement to the caledonian geosyncline (greenlandian) would have been transformed to such a degree as to be unrecognisable (see also reviews of early work in haller 1971, pp. 6–35). up to 1950 no outcrops of the caledonian foreland had been recognised in east greenland. however, the period 1952–1961 saw the discovery and description of foreland areas in kronprins christian land in the north (fränkl 1954, 1955), in dronning louise land (peacock 1956, 1958), and in western gåseland in the south (wenk 1961; see fig. 1). all of these discoveries were incorporated into haller’s major reviews and map compilations of the east greenland caledonides (haller 1970, 1971, 1983; koch & haller 1971). the area around målebjerg (73°27′n) was first mapped by john haller in 1949–1950 as part of a regional investigation of western andrée land (haller 1953): it was a key area for his development of the ‘stockwerke’ models of the caledonian orogeny. on the basis of his studies in andrée land, haller reached wide-ranging conclusions as to the nature of the caledonian orogeny, elaborating on the earlier interpretations of backlund (1930, 1932) and wegmann (1935). thus haller (1953) stated in his english summary (p. 190) that: “the dispute as to the age of the ‘central metamorphic complex’ is, as far as the region of kejser franz josephs fjord is concerned, finally resolved by the present study, which shows that several stratigraphic subdivisions, recognised in the eleonore bay formation, can be traced also in the gneisses, schists and marbles of the ‘central metamorphic complex’. the crystalline rocks of sedimentary origin represent members of the groenlandium, metamorphosed and metasomatically altered during the caledonian orogeny.” with respect to the area around målebjerg, he correlated the distinctive quartzites and dolomites found there with fränkl’s (1951) ‘alpefjord series’, i.e. the lower levels of the eleonore bay ‘formation’ (haller 1953). the ‘groenlandium’ (also speltgrönlandium,groenlandian or greenlandian) is a now obsolete term that originally encompassed all proterozoic sedimentary (and metasedimentary) rocks of north and east greenland (koch 1930). within the caledonides of east greenland it was considered to be made up entirely of the eleonore bay ‘formation’ and equivalents. thus, haller envisaged the sedimentary rocks of the present-day målebjerg window, together with those of the structurally overlying niggli spids and hagar bjerg thrust sheets and franz joseph allochthon, to form parts of a single stratigraphical succession. the gneisses, which he viewed originally as synorogenic granites, formed parts of the mobilised infrastructure (see below). this basic interpretation remained essentially unchanged in haller’s detailed studies of nearby areas (wenk & haller 1953; haller 1955), and was only slightly modified in his later regional descriptions (haller 1970, 1971, 1983; koch & haller 1971). haller’s early observations in the crystalline rocks of the inner fjords and nunatak region (haller 1953, 1955, 1956; wenk & haller 1953), and those of his coworkers, were all interpreted within the context of the ‘stockwerke’ concept. a series of categorical statements in their published descriptions appear to be aimed particularly at countering the interpretations of british geologists. thus they stated: “the archaean basement of the upper algonkian-ordovician series of deposits has hitherto not been found anywhere in central east greenland” (haller 1956, p. 160); “the geologists participating in the investigations agree that the base of the eleonore bay group is not to be found within the central zone of the caledonides” (wenk 1961, p. 8); “the granitic and migmatitic infrastructure, in its present state of preservation, is not older, but younger, than the pre-cambrian sedimentary cover. geus bulletin 6.pmd 10-02-2005, 09:5482 83 the ascent of the granitic solutions, and associated thermal fronts, represented the most important act of the east greenland orogeny” (wenk & haller 1953, p. 32–33). however, within a few years these early interpretations were modified significantly, with recognition that a ‘basement’ to the metasedimentary rocks was in fact recognisable: “inside the caledonian domain, rock units which were originally from the ancient basement, represent substantial ingredients of the fold belt” (haller & kulp 1962, p. 18). the basement gneiss complexes were, nevertheless, considered to have been petrogenetically rejuvenated over large areas, and were assigned a caledonian age. a sketch map by haller (haller & kulp 1962, fig. 3b; haller 1971, fig. 15b) shows that the ‘niggli spids dome’ and ‘gletscherland migmatite complex’ (two of the units of the ‘central metamorphic complex’), were now to be considered caledonian reworked precambrian basement rocks. however, in respect of the nappe-like convolutions of the so called ‘hagar migmatite sheet’, an entirely caledonian origin was still envisaged. in his description of these nappelike migmatite sheets haller 1971 (p. 179) writes: “they are asymmetrical and have considerable overlaps of up to 25 km. regional thrusting is probably of no importance in their formation.” this is essentially a re-statement of the earlier conclusion of wenk & haller (1953, p. 32): “we cannot believe that these structures are due to far-reaching tectonic transport, produced by tangential compression. their mode of occurrence, especially their diapir-like character, indicates that we are here dealing with mobile masses of the infrastructure, which have ascended and intruded into the covering sedimentary series.” many of the essential principles of the ‘stockwerke’ interpretation of the east greenland caledonides were maintained in haller’s later publications (haller 1971, 1983). haller does express regret, however, that the term ‘synorogenic granite’ had been retained on his maps for the central metamorphic complex gneisses (koch & haller 1971), because it was at variance with the modified views given by haller & kulp (1962) and had thus been misunderstood by many workers. the assumption that all metasedimentary rocks in the southern part of the caledonides were parts of the eleonore bay ‘formation’ or ‘group’ (promoted to a ‘supergroup’ by sønderholm & tirsgaard 1993) was retained in haller’s latest publications (1971, 1983). it was not until the 1968–1972 expeditions by the former geological survey of greenland (ggu) to the scoresby sund region that suspicions grew that the widespread high-grade and often migmatitic metasedimentary rocks might be significantly older than the distinctive succession now known as the eleonore bay supergroup. rb-sr whole rock isochrons and u-pb bulk zircon determinations indicated that some granites emplaced into the metasediments were approximately 1000 ma old (hansen et al. 1978; steiger et al. 1979; rex & gledhill 1981). these presumed older metasedimentary rocks, cut by the c. 1000 ma granite suite, were therefore distinguished as the krummedal supracrustal sequence (e.g. henriksen & higgins 1969; higgins 1974, 1988). convincing proof that the krummedal sequence metasedimentary rocks had experienced an early neoproterozoic thermal event (940– 910 ma) not seen in the eleonore bay supergroup had to await the advent of sophisticated modern geochronology, notably ion microprobe age determinations on individual zircon grains (strachan et al. 1995; jepsen & kalsbeek 1998; kalsbeek et al. 2000; watt et al. 2000; leslie & nutman 2000, 2003). in the mid-1970s a number of reconnaissance investigations were carried out in the general kong oscar fjord region. these extended the general conclusions of the 1968–1972 investigations in the scoresby sund region northwards. the widespread mediumto highgrade metasedimentary successions in the kong oscar fjord region were correlated with the krummedal supracrustal sequence (higgins 1988), and rb-sr whole rock isochrons on the underlying gneisses yielded palaeoproterozoic ages (rex & gledhill 1981). two further important observations were made: (1) the distinctive quartzite in the målebjerg window was mapped in 1976 as resting unconformably on the gneissic basement (tage thyrsted in higgins et al. 1981, fig. 8), and (2) during a survey reconnaissance helicopter flight to the eleonore sø region, it was recorded that the base of the ‘slottet quartzite’ at one locality was an unconformity with a basal conglomerate, rather than the supposed thrust. unfortunately, the regional significance of these observations was not appreciated at the time. målebjerg the målebjerg area was a significant location for haller’s ‘stockwerke’ interpretation, in particular for the magnificent exposures of what was described as the ‘zone of detachment’ between the mobile granitic infrastructure and the more rigid metasedimentary superstructure (haller 1971, photograph 47, p. 138; geus bulletin 6.pmd 10-02-2005, 09:5483 84 see also fig. 4). haller mapped the main lithological units in the målebjerg area including the white quartzite now distinguished as the slottet formation, and recognised that they occupied an anticlinal structure (haller 1953, 1970). he observed that the intensity of metamorphism was weaker than elsewhere and that the stratigraphy was well preserved, but he referred this quartzite and nearby marble units to the ‘basal series’ of the eleonore bay ‘group’ (haller 1971, p. 86). haller’s geological and structural maps of the area around målebjerg show several features marked with thrust symbols, some of which have no obvious significance and appear to be photogeological interpretations, while others can be identified with major structures. the most prominent structure on his maps corresponds to the nw–se-trending, ne-dipping thrust that crosses gemmedal (fig. 5), a major feature now identified as the hagar bjerg thrust at the base of the hagar bjerg thrust sheet. haller’s ‘zone of detachment’ in the cliff of målebjerg is not marked with a thrust symbol on his maps, but the folded quartzites of his ‘zone of detachment’ lie immediately beneath the niggli spids thrust of current usage (nst in fig. 4). there is no evidence that haller considered the målebjerg area to be a foreland window or that he suspected the presence of lower palaeozoic sediments. however, he was aware that a “detailed exploration of this key locality is still lacking” (haller 1971, p. 86), and there is reason to believe that such exploration would have been carried out if lauge koch’s expeditions had not been brought to an unexpected close after 1958. as noted above, reconnaissance investigations in the målebjerg area in 1976 led to recognition of an unconformity at the base of the c. 200 m thick quartzite, with a distinctly diverging foliation in the underlying basement gneisses (t. thyrsted in higgins et al. 1981, fig. 8). detailed studies in 1997–1998 by leslie & higgins (1998) established the presence of a significant foreland window (fig. 5). key observations included: (1) the presence of a local diamictite in depressions in the peneplained gneiss surface, now correlatedwith thediamictites of the vendian tillite group; (2) a skolithos-bearing quartzite (the early cambrian slottet formation) unconformably overlying the gneisses and the diamictite; (3) a lower palaeozoic fig. 4. the west face of målebjerg in western andrée land (for location see fig. 5). light coloured folded quartzites (< 200 m thick) of the slottet formation (sf) unconformably overlie grey gneisses (g) that are probably of palaeoproterozoic age. the unconformity is strongly folded. a few metres of grey dolomite (målebjerg formation) occur immediately beneath the niggli spids thrust (nst). overlying units of the niggli spids thrust sheet are dominated by massive mica schists with pale coloured carbonate-rich units (krummedal supracrustal sequence). the summit of målebjerg at right is 1873 m high, about 1500 m above the glacier surface in the foreground. geus bulletin 6.pmd 10-02-2005, 09:5484 85 dolomite unit (målebjerg formation) above the slottet formation quartzite; (4) a major thrust at the top of the målebjerg formation (niggli spids thrust). these observations were confirmed by smith & robertson (1999), who made additional detailed observations and measured sections in the diamictite unit (interpreted as a tillite), and in the slottet and målebjerg formations (smith et al. 2004, this volume). eleonore sø investigation of the western nunatak region between 72° and 75°n was for many years limited to the traverse of the eleonore sø area by hans katz in 1951 (katz 1952, 1953), traverses by eduard wenk and john haller west and south of petermann bjerg in 1951 and 1953 (wenk & haller 1953; haller 1956), and extensive aerial reconnaissance by john haller. cautious statements about what might be present in this vast nunatak area thus amounted to speculation on the basis of very limited ground information. reviewing the possibilities subsequent to the discovery of the caledonian foreland areas in dronning louise land (76°n; peacock 1956, 1958) and in gåseland (70°n; wenk 1961), haller (1971, p. 195–196) suggested that the thrusts around the gåseland window (70°n) and near charcot land (72°n; vogt 1965) had only modest displacements (20 km and < 1 km, respectively). recording that no further outcrops of the foreland were known, he further speculated (haller 1971, p. 218): “considering the structure pattern hitherto obtained from this poorly exposed and little known nunatak region, i would not be surprised if future investigators were able to trace relics of early caledonian overthrust tectonics, on which the present pattern of main folding was then superimposed.” he continues: “however, the main caledonian structures displayed in the wellexplored fjord region are definitely not far travelled; on the contrary, they appear to be autochthonous, initiated and caused by the rise of the migmatite front resulting in a ‘stockwerk’ folded belt.” on the basis of his 1951 traverse of the eleonore sø region, katz had observed that the low grade sedimentary rocks at eleonore sø “are of the same type as those of the eleonore bay formation of the fjord zone” ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ målebjerg gemmedal g erard de g eer g letscher thrust thrust krummedal metasediments (mesoproterozoic) krummedal metasediments (mesoproterozoic) carbonate (krummedal sequence) målebjerg and slottet formations orthogneiss (palaeoproterozoic) orthogneiss (palaeoproterozoic) orthogneiss (palaeoproterozoic) granite 930 ma or caledonian ▲ ▲ ▲ ▲ h ag ar b je rg th ru st sh ee t n ig gl i s pi ds th ru st sh ee t m ål eb je rg fo re la nd w in do w 5 km 73 30' 27 fig. 5. geological map of the målebjerg area, after maps and interpretations of leslie & higgins (1998, 1999). legend below figure illustrates the new thrust terminology. geus bulletin 6.pmd 10-02-2005, 09:5485 86 fig. 6. geological maps of the eleonore sø area, at the same scale but with slightly different topographic bases. a: redrawn after maps and interpretations of katz (1952), haller (1970, 1971) and koch & haller (1971). note legend below figure assigns all sediments and metasedimentary rocks to the eleonore bay ‘group’. b: redrawn after maps and interpretations of leslie & higgins (1998, 1999), using a modern topographic base. legend at top left distinguishes foreland lithologies from allochthonous thrust units. ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ i i i i i i i i i i i i i i i i i i i i i i i i i i ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ 74° 74° 28° 28° arnold escher land arnold escher land harald grieg fjelde harald grieg fjelde slottet slottet wilkins nunatakker j.l. mowinckel land j.l. mowinckel land wilkins nunatakker louise boyd landlouise boyd land a b eleonore sø pa la eo pr ot er oz oi c fo re la nd w in do w a llo ch th on ou s th ru st u ni ts el eo no re b ay ‘g ro up ’ ▲ ▲ palaeogene basalt caledonian granite unconformity caledonian ophiolite (greenstone) caledonian reactivated gneissba sa l s er ie s metasediments argillaceous–arenaceous shale carbonate slottet quartzite fault, tick on downthrow side thrust palaeogene basalt caledonian granite unconformity eleonore bay supergroup krummedal metaseds. (mesoproterozoic) orthogneiss (palaeoproterozoic) målebjerg and slottet formations carbonate and shale metavolcanics granite gneiss ▲ ▲ i i 25 km ham be rg gl ets ch er ham be rg glet sch er geus bulletin 6.pmd 10-02-2005, 09:5486 87 (katz 1953, p. 12). the associated volcanic rocks were interpreted as ophiolitic intrusions and of the same age as the tillites of the fjord zone. while katz correctly depicted a thrust at the west side of harald grieg fjelde (fig. 6a), he considered it to be of devonian age (this thrust corresponds to the major thrust contact on the east side of the window above the målebjerg formation dolomites; fig. 6b). the map in katz (1953) also shows a continuous thrust contact at the base of his ‘slottet quartzite’ (fig. 6a), with the internal structure of the quartzite depicted on his crosssections (katz 1953, tafel 4) as discordant to the ‘thrust’ at the base. it was katz’s interpretation of this prominent quartzite unit as equivalent to the lower part of the eleonore bay ‘formation’ that led him to introduce the basal thrust, because the quartzite lay structurally above the sedimentary rocks he viewed as correlateable with the upper part of the eleonore bay ‘formation’. katz’s cross-section of the eleonore sø region (1953, tafel 4) does indicate tectonic contacts at both margins of the present window, but his preferred interpretation was that the rock units occupied a graben. based on the work of katz, haller initially agreed with katz’s interpretation that the eleonore sø sediments, which from aerial observations he had traced southwards through j.l. mowinckel land to hamberg gletscher, occupied a large post-caledonian graben structure (haller 1956, p. 161). most geologists that have worked in east greenland have come across the widespread and often large erratic blocks of skolithosbearing quartzites, and haller (1971, fig. 48) had plotted observations of these quartzites, and inferred that the source areas lay beneath the inland ice. however, he clearly did not make any link between the skolithos erratic boulders and the ‘slottet quartzite’ of katz (1952, 1953), since he placed this latter unit in his ‘basal series’ of the eleonore bay ‘formation’. summarising the situation at eleonore sø, haller writes of the nonmetamorphic dolomites and quartzites at eleonore sø as being associated with greenschists that katz had interpreted as caledonian ophiolites. he notes that the region in which the outcrops are found is bounded on both sides by tectonic lineaments, and he writes (haller 1971, p. 86–87) “it is open to question whether we are concerned here with parts of the overridden caledonian foreland, similar to the gaaseland ‘window’, or not.” the structures distinguished by haller (1970, 1971) in the vicinity of eleonore sø compared to those mapped by leslie & higgins (1998, 1999) are shown in fig. 6. the only thrust correctly depicted by haller, who here followed the usage of katz (1952), is the east-dipping structure on the east side of the window at harald grieg fjelde. at this locality, high-grade metasedimentary rocks in the hanging wall lie structurally above low-grade carbonates and quartzites in the foot wall. in view of the difficulties of access to the region in the 1950s, and the lack of isotopic age determinations, it is not surprising that the ages attributed to the rock units in the eleonore sø region by katz and haller have since proved to be incorrect. however, the thrust that marks the west side of the window is depicted on haller’s maps as a major normal fault, whereas the other thrusts shown on haller’s interpretation (fig. 6a) correspond to the unconformity at the base of the ‘slottet quartzite’, the presentday slottet formation. it is, perhaps, surprising that katz did not apparently examine the base of the quartzite unit, which is an obvious unconformity in the field, and often has a basal conglomerate. his erroneous interpretation of this unconformity as a thrust contact does give the impression of an arched thrust on haller’s structural maps (e.g. haller 1971, fig. 58), and has been taken by some authors as evidence that haller ‘discovered’ the window (cf. hartz et al. 2001). however, as noted above, haller (1971) considered it “open to question”. only a few geologists have visited the eleonore sø region since katz’s 1951 visit, but it was not until the survey’s 1997–1998 regional mapping expedition that detailed field studies led to the regional delineation of the basal unconformity of the ‘slottet quartzite’. the authors of this article retraced one of katz’s traverses eastwards towards harald grieg fjelde, and observed that the basal contact of the quartzite was in fact conformable to the bedding within the quartzite rather than discordant as in katz’s profiles (katz 1952). the quartzites were observed to contain well-preserved sedimentary structures, and the first in situ finds of skolithos were found just west of the thrust at harald grieg fjelde. the unconformity surface, subsequently studied at several localities, proved to be a clean undisturbed contact (fig. 7) with a basal conglomerate up to 1.5 m thick often present. in nearly every section of the quartzites examined, in situ long skolithos burrows were observed, and these demonstrate that the sequence now termed the slottet formation (smith et al. 2004, this volume) is of lower cambrian age (crimes 1992). a cambrian–ordovician age can therefore be assumed for the thin dolomite sequence of the målebjerg formation, which conformably overlies geus bulletin 6.pmd 10-02-2005, 09:5487 88 the slottet formation and immediately underlies the thrust (leslie & higgins 1998, 1999; smith et al. 2004, this volume). on the west side of the window, the west-dipping thrust contact is well exposed and associated with thick developments of mylonites. a clear definition of the entire thrust system bordering the window was established for the first time during the 1997–1998 mapping (fig. 6b). finds of low grade volcanic rocks in eastern j.l. mowinckel land have extended the known distribution of their occurrence considerably, such that the eastern marginal thrust of the window can now be placed parallel to the glacier to the east of j.l. mowinckel land. exposures of characteristic rock types (low-grade volcanic rocks and skolithos-bearing quartzites) observed as far north as nunataks at 74°25′n (fig. 6b) show that the nne– ssw extent of the eleonore sø window is at least 125 km. the volcanic succession (pillow lavas and tuffs) and associated sedimentary rocks (thick dolomites and dolomite breccias, sandstones and shales) that are unconformably overlain by the slottet formation quartzites are intruded locally by quartz porphyry bodies, dated by shrimp analyses of zircon to c. 1950 ma (f. kalsbeek, personal communication 2000). the eleonore sø volcano-sedimentary rocks are thus palaeoproterozoic or older in age, and can be broadly compared with the volcano-sedimentary rocks of the charcot land window at c. 72°n and the foreland exposures of the hamberg gletscher complex (volcanic rocks and associated gabbros) at c. 73°n (fig. 1; higgins et al. 2001). discussion the exceptional exposures in the extensive fjord system and nunataks of the kong oscar fjord region (72°– 75°n), and the long series of geological expeditions led by lauge koch (1926–1958), have deservedly led to recognition of the east greenland caledonides as a spectacular example of an orogenic belt. this is in large part a tribute to the superb compilations of data presented by john haller (haller 1970, 1971; koch & haller 1971). development of new models for the east greenland caledonides, to replace the ‘stockwerke’ concept, has been a gradual process extending over a period of some 30 years (1968–1998), during which the entire 1300 km length of the orogen has been remapped as part of the survey’s regional 1:500 000 mapping project. during this extended period of refig. 7. the nunatak slottet in the eleonore sø window, looking northwards. the white (lower) and dark (upper) quartzites of the lower cambrian slottet formation (sf: 350 m thick) rest unconformably on dark coloured palaeoproterozoic clastic sediments of the eleonore sø volcano-sedimentary complex (es). the highest summit of slottet is 1933 m high, about 600 m above the glacier surface. geus bulletin 6.pmd 10-02-2005, 09:5488 89 search, most of the assumptions built into haller’s (1970, 1971) ‘stockwerke’ concept of an in situ caledonian orogenic belt have been queried or refuted, and the new interpretations confirmed by increasingly sophisticated isotopic age determinations. prior to the survey’s 1997–98 expeditions the existence of far-travelled thrust sheets had not been demonstrated in the kong oscar fjord region. with discovery of the eleonore sø and målebjerg windows, and distinction of a thrust pile with hundreds of kilometres of west-north-west thrust displacement (henriksen 1998, 1999; leslie & higgins 1998, 1999; elvevold et al. 2000; higgins et al. 2004), the ‘stockwerke’ concept of in situ caledonian orogenesis can finally be pronounced dead and laid to rest. restoration of the thrust sheets to their approximate original locations implies that the focus of caledonian orogenesis, i.e. the collision of laurentia with baltica, took place several hundred kilometres east-south-east of the orogenic belt now preserved onshore in east greenland. wordie (1930) and parkinson & whittard (1931) were, in fact, partly correct when they compared the crystalline gneisses of the inner fjord region of east greenland to the archaean lewisian gneisses of scotland. the former have yielded archaean and proterozoic protolith ages, with archaean gneiss complexes extending throughout the inner part of the scoresby sund region and northwards to southern suess land (72°50′n). farther north the orthogneisses have yielded palaeoproterozoic protolith ages which relate to an important episode of regional palaeoproterozoic crustformation well documented throughout the northern half of the east greenland caledonides (kalsbeek et al. 1993, 1999). it follows that haller (1953) was wrong in attributing the formation of the orthogneisses to rising fronts of caledonian migmatisation that transformed metasedimentary rocks of the eleonore bay ‘formation’. haller’s later re-interpretation of these gneisses as caledonian reworked basement rocks was close to the present-day interpretation (haller & kulp 1962; haller 1971). john haller was of swiss nationality, educated in switzerland, and obviously familiar with the major thrusts and fold nappes of the alpine orogenic belt. his earliest 1949–1951 studies in east greenland were in andrée land (haller 1953), and his main conclusions were presented as a confirmation and elaboration of the interpretations of h.g. backlund and c.e. wegmann. he was already committed to the idea of widespread transformation of a single metasedimentary succession (eleonore bay ‘group’) by the vertical rise of mobile migmatitic bodies, a view that was developed during his subsequent field work, and elegantly presented as the ‘stockwerke’ concept (haller 1970, 1971). haller’s wide-ranging observations on the ground and from the air, and katz’s observations around eleonore sø, were all interpreted within the context of the basic ‘stockwerke’ model. thus the palaeoproterozoic volcano-sedimentary succession of the eleonore sø window and the thin lower palaeozoic successions of the målebjerg and eleonore sø windows were referred to the eleonore bay ‘group’. an unconformity at the base of the ‘slottet quartzite’ in the eleonore sø window, presumably not examined very closely, was interpreted as a major thrust in order to force the stratigraphy to fit into the model. the displacements on the major thrusts that were recognised were grossly underestimated, perhaps in order not to upset the assumption that “the main caledonian structures displayed in the well-explored fjord region are definitely not far travelled; on the contrary, they appear to be autochthonous” (haller 1971, p. 218). while the ‘stockwerke’ model of intense in situ granitisation is no longer tenable, the caledonian orogeny in east greenland was certainly not the ‘superficial’ orogeny envisaged by the early british geologists. the precambrian orthogneiss complexes, together with the overlying metasedimentary successions, have experienced high-grade caledonian metamorphism and intense reworking during the regional caledonian compressive deformation that produced major westward propagating thrust sheets. caledonian granites generated by melting of mesoproterozoic sediments are widespread in the hagar bjerg thrust sheet, but absent in the lower niggli spids thrust sheet. the dominant fabric in the archaean and palaeoproterozoic orthogneisses of the thrust sheets is today interpreted in many areas to be essentially caledonian, which as a concept is not greatly different from the ‘caledonian petrogenetic rejuvenation’ envisaged by haller & kulp (1962, p. 18). however, despite caledonian reworking, the orthogneisses of the crystalline complexes still yield archaean and palaeoproterozoic protolith ages, and in low strain areas relicts of the precambrian foliation cut by discordant amphibolite dykes are preserved (higgins et al. 1981, p. 37–38). the assumption that all metasedimentary rocks in the southern half of the east greenland caledonides were variably transformed parts of the eleonore bay ‘group’, was not questioned until ggu’s work in the scoresby sund region in 1968–1972. although then geus bulletin 6.pmd 10-02-2005, 09:5489 90 based on imprecise rb-sr and u-pb ages (rex & gledhill 1974; hansen et al. 1978; steiger et al. 1979), ggu’s investigations led to distinction of two sedimentary successions (henriksen & higgins 1969, 1976; higgins 1974, 1988). the widespread high-grade metasedimentary rocks that hosted c. 1000 ma augen granites were ascribed to the krummedal supracrustal sequence, whereas the high-grade to non-metamorphic eleonore bay supergroup appeared to be affected only by caledonian metamorphism and deformation. this viewpoint did not go unchallenged, and diverging interpretations have continued to be expressed (peucat et al. 1985; hartz & andresen 1995; andresen & hartz 1998; hartz et al. 2000). recent ion microprobe zircon studies have now confirmed the widespread distribution of a distinctive 940–910 ma granite suite hosted by the high-grade, commonly migmatitic, krummedal sequence of the hagar bjerg thrust sheet (jepsen & kalsbeek 1998; kalsbeek et al. 2000; leslie & nutman 2000; watt et al. 2000; watt & thrane 2001). a later granite suite, hosted by both the krummedal sequence and the lowest part of the eleonore bay supergroup, is caledonian in age (rex & gledhill 1981; hartz et al. 2001; kalsbeek et al. 2001a, b; white et al. 2002). the most dramatic revelation of the recent survey mapping is that the 18.5 km thick neoproterozoic– ordovician succession preserved in the franz joseph allochthon of the hagar bjerg thrust sheet structurally overlies a partly equivalent < 400 m thick sequence preserved in the målebjerg window (fig. 4). higgins et al. (2001, fig. 8) demonstrated the similarities of the restricted foreland succession of the målebjerg window with that in the eleonore sø window and other foreland areas preserved along the western margin of the east greenland caledonides. it follows that the allochthonous and very thick eleonore bay supergroup – tillite group – kong oscar fjord group succession must have been laid down in a completely different sedimentary environment a substantial distance to the east of the restricted sequence deposited on the foreland craton. the succession preserved in the franz joseph allochthon of east greenland exhibits broad similarities with the major neoproterozoic – lower palaeozoic sedimentary successions of svalbard, nw scotland and newfoundland that were deposited along the western passive margin of the iapetus ocean (e.g. swett & smit 1972; soper 1994). in east greenland the preserved remnants of this basin were displaced at least 200 km, and possibly as much as 400 km, westnorth-west across the laurentian margin to structurally overlie their thin foreland equivalents, a caledonian shortening across the orogenic belt estimated at 40–60% (higgins & leslie 2000; higgins et al. 2001, 2004). as noted above, restoration of the thrust sheets to their approximate original locations implies that the collision of laurentia with baltica took place several hundred kilometres east-south-east of the orogenic belt now preserved onshore in east greenland. the models of the caledonian orogen presented by hartz & andresen (1995) and andresen et al. (1998), which invoked upward and lateral movement of light, low viscosity, lower crustal material towards the region of maximum crustal extension, a process compared to haller’s ‘stockwerke’ concept, neglect the significance of caledonian thrusting. following the survey’s demonstration of the existence of the foreland windows and the presence of major thrusts at caledonian symposiums held in copenhagen (frederiksen & thrane 1998, 1999), a considerably revised model for the orogen was presented by hartz et al. (2001). while the thrust terminology employed by hartz et al. (2001) has similarities with that of elvevold et al. (2000), there are many differences in interpretation. some boundaries on their map (hartz et al. 2001, fig.1) appear to have been adopted from koch & haller’s (1971) obsolete maps, and, for example, the west-dipping thrust of the eleonore sø window is incorrectly indicated as an east-dipping extensional fault (cf. fig. 6). john haller’s contributions to the understanding of the east greenland caledonides are considerable (see e.g. henriksen & higgins 1993; schwarzenbach 1993). however, although he did not discover the målebjerg and eleonore sø windows or identify the lower palaeozoic rock units, many of his observations, in retrospect, support such an interpretation. unfortunately, haller’s emphasis on the autochthonous in situ origin of the crystalline complexes led him to deny that significant thrusting was involved in the central fjord zone, and to underestimate displacements on the thrusts that were observed in gåseland, near charcot land, around eleonore sø, and around målebjerg in andrée land. acknowledgements the helpful comments of the two reviewers, brian chadwick and arild andresen, are gratefully acknowledged, and have resulted in correction of a number of errors and improvements to the original text. geus bulletin 6.pmd 10-02-2005, 09:5490 91 references andresen, a. & hartz, e.h. 1998: basement–cover relationships and orogenic evolution in the central east greenland caledonides. gff 120, 191–198. stockholm: geological society of sweden. andresen, a., hartz, e. & vold, j. 1998: a late orogenic extensional origin for the infrastructural gneiss domes of the east greenland caledonides (72°–74°n). tectonophysics 285, 353–369. backlund, h.g. 1930: contributions to the geology of northeast greenland. meddelelser om grønland 74(11), 207–296. backlund, h.g. 1932: das alter des ‘metamorphen komplexes’ von franz josef fjord in ost-grönland. meddelelser om grønland 87(4), 119 pp. best, m.g. 2003: igneous and metamorphic petrology (2nd edit.), 729 pp. oxford: blackwell publishing. cowie, j.w. & adams, p.j. 1957: the geology of the cambroordovician rocks of central east greenland. part 1. stratigraphy and structure. meddelelser om grønland 153(1), 193 pp. crimes, t.p. 1992: the record of trace fossils across the proterozoic–cambrian boundary. in: lipps, j.h. & signor, p.w. 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(eds): geology of east greenland 72°–75°n, mainly caledonian: preliminary reports from the 1998 expedition. danmarks og grønlands geologiske undersøgelse rapport 1999/19, 11–26. leslie, a.g. & nutman, a.p. 2000: episodic tectono-thermal activity in the southern part of the east greenland caledonides. geology of greenland survey bulletin 186, 42–49. leslie, a.g. & nutman, a.p. 2003: evidence for neoproterozoic orogenesis and early high temperature scandian deformation events in the southern east greenland caledonides. geological magazine 140, 309–333. parkinson, m.m.l. & whittard, w.f. 1931: the geological work of the cambridge expedition to east greenland in 1929. quarterly journal of the geological society (london) 87, 650–674. peacock, j.d. 1956: the geology of dronning louise land, n.e. greenland. meddelelser om grønland 137(7), 38 pp. peacock, j.d. 1958: some investigations into the geology and petrography of dronning louise land, n.e. greenland. meddelelser om grønland 157(4), 139 pp. peucat, j.j., tisserant, d., caby, r. & clauer, n. 1985: resistance of zircons to u-pb resetting in a prograde metamorphic sequence of caledonian age in east greenland. canadian journal of earth sciences 22, 330–338. rex, d.c. & gledhill, a.r. 1974: reconnaissance geochronology of the infracrustal rocks of flyverfjord, scoresby sund, east greenland. bulletin of the geological society of denmark 23, 49–54. rex, d.c. & gledhill, a.r. 1981: isotopic studies in the east greenland caledonides (72°–74°n) – precambrian and caledonian ages. rapport grønlands geologiske undersøgelse 104, 47– 72. schwarzenbach, f.h. 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leth 1996, 1998). knowledge on the sedimentary conditions and processes along the entire west coast of jutland has, however, significantly increased as a result of 2000 km of newly acquired high-resolution seismic and side-scan sonar data, supplemented by about 100 vibrocores. these data were collected by the geological survey of denmark and greenland (geus) during joint projects with the danish coastal authority between 1998 and 2001 (leth et al. 1999; larsen & leth 2001). the coastal zone off west jutland displays a highly dynamic environment, where sediment transport is governed by strong tidal and wave-induced currents. the net wavegenerated current is south going, while the coastal current has a net direction towards the north (knudsen et al. 2002). the direction of the net littoral drift is southward from the outlet of limfjorden to blåvands huk, with net erosion north of nymindegab and aggregation to the south; the offshore part of this depositional system has recently been studied (larsen & leth 2001). seabed geology the present coast between nymindegab and limfjorden is a highly erosive, wave-dominated, high-energy barrier coast subdivided by glacial headlands. the natural retreat of the coast profile is in the order of 2–5 m/year. mapping of the geological substratum reveals a complex range of lithologies, of which five types are dominant: (1) weichselian glacial till, (2) saalian till, (3) weichselian outwash deposits, (4) eemian marine deposits and (5) miocene sediments (fig. 2). the development of the coast and sand transport has been closely linked to the subsurface geology since the onset of the holocene transgression and up to the present (leth 1996, 1998; anthony et al. in press). distribution and transport of mobile sand north of nymindegab the thickness and regional distribution of the upper layer of fineto medium-grained marine sand have been mapped (fig. 3). the accumulations can be considered as positive morphological forms. twelve samples from the sand have yielded dates from 150 to 835 years b.p. (calibrated c-14 radiocarbon ages). the sand unit is therefore regarded as mobile sand representing the recent to subrecent hydrographic regime. a regional erosional unconsediment distribution and transport in the shallow coastal waters along the west coast of denmark jørgen o. leth, birger larsen and dennis anthony fig. 1. location map of the studied offshore coastal zone covering approximately 2300 km2. the two boxes (red) mark the key areas for the study. the depth interval is 5 m. geological survey of denmark and greenland bulletin 4, 41–44 (2004) © geus, 2004 42 fig. 2. map showing the seabed geology below the mobile sand layer. fig. 3. map showing the thickness and distribution of the mobile sand layer. formity underlying the sand layer suggests that the hydrodynamic energy level in the coastal zone may have increased within the past millennium. as the unit is less than 0.5 m thick in most of the study area, the sand available for sediment transport in the area is obviously limited. however, the mobile sand occurs as ne–sw-trending shoreface-connected ridges close to the coast between nymindegab and bovbjerg (anthony & leth 2002) that extend into large sand-bodies farther offshore. the formation of two large sand accumulation areas, 15 and 30 km long and with a nnw–sse trend, is not clearly understood (fig. 3). their existence is most likely due to a combination of a hydrographically controlled sediment transport pattern and the presence on the nearby seabed of sandy outwash deposits. due to the young age of the deposits and the shape of present bedforms, it is suggested that the jutland coastal current is the major agent responsible for the reworking and redistribution. the overall oblique form of the sand relative to the coast is explained by the fact that the offshore nnw-directed tidal flood current is stronger than the tidal ebb current. that the northern sand body is much larger than the southern seems to reflect the general northwards decrease of the tidal strength, leading to a decrease in tidal current sediment transport capacity towards the north, and thus net deposition. it is probable that a sediment transport system interacting between the study area and the adjacent area further west exists, but to test this more research is needed. the prograding coast the coast south of nymindegab is aggrading at about 0.5–2 m/year. over a distance of 25 km it has prograded some 3 km westwards in the form of a beach ridge system during the last 3000 years (nielsen et al. 1995; larsen & andersen in press). the development of the back-barrier area from 8000 to 3000 years b.p. is not well known because of the 5–10 m thick cover of aeolian sand. most of the deposition in this area took place in lagoons, like that of the present day ho bugt, in lakes dammed between the littoral deposits and the saalian landscape, or in the form of large dunes and cover sands. the south-east-trending peninsula skallingen is a very young (after a.d. 1600) analogue to a barrier island. the westwardprotruding bank of inner horns rev belongs to the same system. blåvands huk forms the centre of a major sand accumulation area some 25 km long, 5–15 km wide and 15–25 m thick, with a total volume of about 6 km3. this holocene spit complex is built out onto an erosional, very flat platform at about 20 m below present sea level, west of the old saalian glacial landscape. the whole complex has accumulated steadily throughout the past 8000 years, and is still very active. inner horns rev is a 6 km wide and 20 km long bank protruding westwards from blåvands huk (fig. 4). new investigations indicate that it is highly dynamic, with active sand accumulation particularly on the slopes. the direction of progradation suggests longshore sediment transport (fig. 4). c-14 datings suggest that it has prograded some 3.5 km westwards during the last 800 years. the shape of the blåvands huk spit complex is due to a combination of wave-dominated longshore sediment transport from the north, tidal influence and the partial shelter from the offshore banks at horns rev. according to recent estimates by the danish coastal authority (kystdirektoratet 2001), some 2.3 million m3 of sediments are supplied annually to the system from the north. a second accumulation system is located at the outer horns rev to the south-west (fig. 4). this is separated from the inner horns rev by the channel slugen. the base of horns rev is an erosional unconformity on top of eemian marine silt in the central part, and on saalian glacial deposits further to the west. a wide valley has been cut into eemian deposits and subsequently filled with weichselian glaciofluvial deposits, with remnants of small lower holocene bogs on the top (fig. 5). from the present study it is now clear that horns rev consists of holocene sand with gravel deposited after the sea transgressed the area in early holocene time, about 8500 years b.p., and is not a saalian terminal moraine as has been previously proposed. precisely how the outer horns rev has formed is not clear, but all sediment structures suggest westerly sediment transport. at southern horns rev the lower 43 fig. 4. sediment transport directions for the holocene marine sand determined from direction of sediment structures, distribution of sediment thickness and modelling of recent sediment transport by the danish hydraulic institute. a: section shown in fig. 5. 44 part is a complex of gravel-rich spits that have grown towards the east-north-east. prograding reflectors and sand waves also suggest a net eastward aggradation. this is in agreement with the annual average direction and magnitude of recent sediment transport, as modelled by the danish hydraulic institute (dhi). on a day-to-day basis the sediment transport is very variable, in response to the strong tidal currents and the breaking waves. sediment transport from a westsouth-westerly direction is also indicated by the direction of progradation and the now buried spits in the up to 15 m thick holocene successions further to the north. the net direction of sediment transport based on this evidence is shown in fig. 4. the blåvands huk – horns rev area has been a major depocenter for sediments transported southwards along the west coast, as well as material presumably eroded from the floor of the north sea, in spite of the very exposed setting. references anthony, d. & leth, j.o. 2002: large-scale bedforms, sediment distribution and sand mobility in the eastern north sea off the danish west coast. marine geology 182, 247–263. anthony, d., leth, j.o., konradi, p. & andersen, l.t. in press: shallow seabed geology and paleogeography off the danish north sea coast. boreas. knudsen, s.b., laustrup, c., madsen, h.t. & christensen, e.d. 2002: sediment transport in the outer part of the coastal profile, 1–13. 28th international conference on coastal engineering, cardiff, wales. kystdirektoratet 2001: kystdirektoratets program for undersøgelser og udvikling, 1998–2001, slutrapport, 36 pp. lemvig: kystdirektoratets kysttekniske afdeling. larsen, b. & andersen, l.t. in press: late quaternary stratigraphy and morphogenesis in the eastern north sea: horns rev and fanoe bay area and its relation to onshore geology. netherlands journal of geosciences. larsen, b. & leth, j.o. 2001: regionalgeologisk tolkning og en samlet vurdering af aflejringsforholdene i området mellem nymindegab og horns rev. danmarks og grønlands geologiske undersøgelse rapport 2001/96, 83 pp. leth, j.o. 1996: late quaternary geological development of the jutland bank and the initiation of the jutland current, ne north sea. geological survey of norway bulletin 430, 25–34. leth, j.o. 1998: late quaternary geology and recent sedimentary processes of the jutland bank region, ne north sea, 173 pp. unpublished ph.d. thesis, university of aarhus, denmark. leth, j.o., anthony, d., andersen, l.t. & jensen, j.b. 1999: geologisk kortlægning af vestkysten. regionalgeologisk tolkning af kystzonen mellem lodbjerg og nymindegab. danmarks og grønlands geologiske undersøgelse rapport 1999/75, 29 pp. nielsen, s.t., andreasen, f. & clemmensen, l.b. 1995: the middle and late holocene barrier spit system at vejers, denmark: structure and development. bulletin of the geological society of denmark 42, 105–119. authors’ addresses j.o.l. & b.l., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: jol@geus.dk d.a., royal danish administration of navigation and hydrography, overgaden o. vandet 62 b, p.o. box 1919, dk-1023 copenhagen k, denmark. fig. 5. geological cross-section of outer horns rev. location shown on fig. 4. the numbers above each borehole indicates displacement relative to cross-section. << /ascii85encodepages false /allowtransparency false /autopositionepsfiles true /autorotatepages /all /binding /left /calgrayprofile (dot gain 20%) /calrgbprofile (srgb iec61966-2.1) /calcmykprofile 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<feff4f7f752890194e9b8a2d5b9a5efa7acb76840020005000440046002065874ef65305542b8f039ad876845f7150cf89e367905ea6ff0c4fbf65bc63d066075217537054c18cea3002005000440046002065874ef653ef4ee54f7f75280020004100630072006f0062006100740020548c002000520065006100640065007200200035002e0030002053ca66f465b07248672c4f86958b555f3002> >> >> setdistillerparams << /hwresolution [2400 2400] /pagesize [595.000 842.000] >> setpagedevice geological survey of denmark and greenland bulletin 13, 2007, 61-64 61 the baltic sea is one of the largest brackish water bodies in the world (segerstråle 1957) with a number of basins varying from almost fresh water in the northern part of the bothnian bay via the more brackish conditions in the southern part to the saline waters of the kattegat. the baltic sea is subject to severe environmental degradation caused by commercial and leisure activities, including fisheries, dredging, tourism, coa s t al development and land-based pollution sources. this causes severe pressures on vulnerable marine habitats and natural re sources, and a tool for aiding marine management is therefore strongly needed. the marine landscape concept presented by roff &taylor (2000) is based on the use of available broad-scale geological, physical and hydrographical data to prepare ecologically meaningful maps for areas with little or no biological information. the concept, which was elaborated by day & roff (2000) was applied in uk waters (connor et al. 2006) before it was adopted by the balance project described here. the aim of developing marine landscape maps is to characterise the marine environment of the baltic sea region (the baltic sea together with the kattegat) using geophysical and hydrographical parameters. such maps can be applied, for example, to an assessment of the baltic-wide network of marine protected areas, and thus provide a sustainable ecosystem-based approach to the protection of the marine environment from human activities, and contribute to the conservation of marine biodiversity. the balance project is based on transnational and cross-sectoral co-operation with participants from nine countries surrounding the baltic sea as well as norway (fig. 1), and is partially financed by the european union through the bsr interreg iiib programme. data collation and harmonisation one of the most challenging aspects of marine landscape map production is collating and harmonising data sets from different sources and with different formats. the data sets include: bathymetry, seabed sediment types, the photic zone, ice cover, halocline depth, temperature, current velocity and bottom salinity. the data sets provided by the individual partners in the balance project were analysed in detail before merging in a gis platform to produce the final benthic marine landscape map (al-hamdani & reker in press). the data sets were obtained through a combination of field measurements and modelling. other data were considered, but not acquired for the entire area; these include oxygen depletion, stratification, wave exposure and pycnocline depth. development of marine landscape maps the uniqueness of the baltic sea region originates in part from its salinity distribution. a stable salinity gradient and stratification is observed in both vertical and horizontal dimensions. this results in wide biogeographic variation as the water salinity changes from marine in skagerrak to nearly fresh waters in the bothnian bay. thus a wide variety of complex marine landscapes reflecting the complexity of the in situ regimes of physical factors is expected. the physical factors chosen for defining the marine landscapes in the baltic sea development of marine landscape maps for the baltic sea and the kattegat using geophysical and hydrographical parameters zyad k. al-hamdani, johnny reker, jørgen o. leth, anu reijonen, aarno t. kotilainen and grete e. dinesen © geus, 2007. geological survey of denmark and greenland bulletin 13, 61–64. available at: www.geus.dk/publications/bull fig. 1. bathymetry of the baltic sea, the kattegat and the skagerrak, the working area of the balance project. data source: geological survey of denmark and greenland, geological survey of sweden and geo logical survey of finland region are considered important for structuring the distribution of major biological assemblages in the region. three physical parameters were adopted in this work to produce the benthic marine landscape map: sediment type, the photic zone, and bottom salinity. the best way to produce a benthic marine landscape map from a number of different sources is by using raster map algebra in a gis platform. this method allows the combining of several different parameters stored in separate layers into a single map layer. one of the major tasks in the production of marine landscape maps for the baltic sea region was to split the chosen environmental data sets in such a way as to produce ecologically relevant classes. the justification for the classification of each data layer is explained separately below, bearing in mind that this is the first attempt to classify such features in the baltic sea; future amendments are thus to be expected. seabed sediment types the data sets for this layer were gathered from different governmental and research institutes of the baltic sea countries and norway. the existing data are abundant and very diverse and have been acquired using different field techniques during the past several decades. seabed sediment maps from offshore and coastal areas exist at a wide range of scales from local (1:20 000) to regional (up to 1:1 000 000). termi nology and classifications vary as well, since the nine circum-baltic nations together with norway have interpreted their own data according to different national classification schemes. national seabed sediment classification categories needed to be harmonised in order to produce the regional map for seabed sediment types (fig. 2). the resulting classification scheme consists of five sediment classes, which can be extracted from existing data. these sediment classes are: 1. hard bottom, including bedrock (crystalline and sedimentary) and bedrock covered with boulders. 2. hard bottom composite, including complex, patchy hard surface and coarse sand (sometimes also clay to boulders). 3. sand, including fine to coarse sand (with gravel exposures). 4. hard clay, sometimes/often/possibly exposed or covered with a thin layer of sand/gravel. 5. mud, including gyttja-clay to gyttja-silt. the photic zone from an ecological point of view, available light is one of the primary physical factors influencing and structuring the biological communities in the marine environment, as it is the driving force behind primary production by providing energy for photosynthesis. the depth of the euphotic zone is traditionally defined as the depth where 1% of the surface irradiance (as measured just below the water surface) is available for photosynthesis. this value was calculated by multiplying the actual measured secchi depths by a factor of 1.9 (a. erichsen, personal communication 2006). based on the irradiation depth, the photic zone was split into two intervals: the euphotic zone and the non-photic zone. these two zones 62 fig. 2. map of the seabed sediments. data source: geological survey of denmark and greenland, geological survey of sweden and geological survey of finland. fig. 3. map of the photic zones. the original modelled data set is a 620 m grid of the average secchi depth measured from march to the end of november for the period from 1980 to 1998. data source: danish hydraulic institute. reflect the significant ecological difference between the shallow-water environment where primary production takes place, and the deeper waters where species and biomass are dominated by fauna and bacteria (fig. 3). bottom salinity salinity is one of the primary physical factors structuring the distribution of species within the kattegat and the baltic sea, varying from almost fresh water in the bothnian bay to normal marine waters in skagerrak (fig. 4). the classification of the salinity data set and justification of the classification are presented in table 1. the three layers, sediment, photic zone, and bottom salinity were combined using the spatial analysis tool in the gis program resulting in the benthic marine landscape map of the baltic sea region shown in fig. 5. this map shows the distribution of 60 marine landscape types each representing different physical conditions at the seabed. the marine land s cape map was further analysed with a statistical tool in gis program to extract the diversity in marine landscape distribution (fig. 6). application of marine landscape maps the marine landscape maps developed for the baltic sea region are a first approach to a broad-scale, physical characterisation of the marine environment of the baltic sea and the kattegat. it will be used throughout the balance project to assess the representativity of the network of marine protected areas (mpas) within the baltic sea region. this assessment will identify whether some marine landscape types are missing or over-represented within the existing mpa network and thus help inform environmental managers whether the existing network is protecting and representing the marine diversity of the baltic region. the landscape map can also be used to show the complexity of the marine environment within a certain area (fig. 5) to enhance future management and protection of the marine ecosystem. in addition to a continuous validation process and confidence rating of the data layers and maps there are many 63 fig. 4. map of the bottom salinity. the original modelled data set is a 7.5 km grid horizontal resolution and 20 vertical layers. the model is based on monthly, averaged values from august 2003 over a period of one year. data source: national environmental research institute, denmark. future challenges for the marine landscape maps. these will include adapting them as a tool for various eu directives implementation, using the maps as (1) a strategic tool for planning future field surveys for mapping (e.g. natura 2000 habitats in the eu habitats directive, (2) a physical characterisation of the marine environment in the proposed marine strategy directive, or (3) part of the typologies in the eu water framework directive. the approach described here should, of course, be improved and adapted to this legislative framework, including more physical layers depending on end-user requirements. similarly, development of pelagic marine landscape maps could be combined with data on e.g. commercial fish species or marine mammals, to provide valuable information to improve management of the marine environment. in conclusion, marine landscape mapping in the baltic sea region is just beginning and although much work has been put into this first step, there are still many challenges ahead. these include access to existing data from the entire baltic region, assigning confidence ratings to the map, improving data layers classification and most importantly, providing good practice examples on how this characterisation can be applied in implementing eu legislation and planning. acknowledgements data, analysis and stimulating discussions for the production of the marine landscape maps were provided by david connors (joint nature conservation committee uk), karsten dahl, johan söderkvist and jørgen bendtsen (national environmental research institute, denmark), jesper h. andersen and anders erichsen, (danish hydraulic institute, water environment & health), lisbeth tougaard (geological survey of denmark and greenland) and daria ryabchuk (all-russian geological institute, st. petersburg). references al-hamdani, z.k. & reker, j. (eds) in press: towards marine landscapes in the baltic sea. copenhagen: geological survey of denmark and greenland. connor, d.w., golding, n., robinson, p., todd, d. & verling, e. 2006: ukseamap: the mapping of marine seabed and water column features of uk seas, 104 pp. peterborough: joint nature conservation committee. day, j.c. & roff, j.c. 2000: planning for representative marine protected areas: a framework for canada’s oceans, 147 pp. toronto: world wildlife fund canada. roff, j.c. &. taylor, m.e. 2000: viewpoint. national frameworks for marine conservation – hierarchical geophysical approach. aquatic conservation. marine and freshwater ecosystems 10, 209–223. segerstråle, s.g. 1957: the baltic sea. in: hedgpeth, j.h. (ed.): treatise on marine ecology and paleocology, 1. ecology. geological society of america memoir 67, 751–800. 64 authors’ addresses z.k.a. & j.o.l., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: azk@geus.dk a.t.k. & a.r., geological survey of finland, p.o. box 96, fin-02151 espoo, finland. j.r. & g.e.d., danish forest and nature agency, haraldsgade 53, dk-2100 copenhagen ø, denmark. fig. 5. map showing the 60 marine landscape types recognised in the baltic sea, the kattegat and the skagerrak. for further details see alhamdani & reker (in press). fig. 6. map showing diversity of marine landscape types. geological survey of denmark and greenland bulletin 35, 2016, 71-74 71© 2016 geus. geological survey of denmark and greenland bulletin 35, 71–74. open access: www.geus.dk/publications/bull in recent years, the greenland ice sheet has been losing mass at an average rate of 262 ± 21 gt yr–1 (2007–2011; andersen et al. 2015). part of this mass loss was due to increases in melt, reducing the surface mass budget (enderlin et al. 2014). also, the acceleration of many marine-terminating outlet glaciers increased the dynamic mass loss (rignot et al. 2008). both mass-loss mechanisms are linked to recent increases in atmospheric and oceanic temperatures (dutton et al. 2015). for instance, in summer 2012 greenland experienced exceptionally warm atmospheric conditions, causing nearly the entire ice-sheet surface to melt for two periods of several days (nghiem et al. 2012) and contributing to the largest annual ice-sheet mass loss on record (khan et al. 2015). this is in contrast to a return to more average conditions in 2015 (tedesco et al. in press). in 2007 the programme for monitoring of the greenland ice sheet (promice) was initiated to monitor both the surface mass budget and dynamic contributions to mass change. for the monitoring, c. 20 automatic weather stations were distributed over eight regions of the greenland ice sheet (fig. 1), primarily in the ablation area where surface melting is most prominent (van as et al. 2011). these stations record a suite of meteorological and radiative variables that allow for surface-energy budget closure, and reveal the relative importance of the different energy fluxes contributing to melting. each station also monitors ablation by sonic height rangers, pressure transducers and ablation stakes (fausto et al. 2012). table 1 shows that the 2015 melt season yielded ablation totals below the promice average (i.e. reduced surface mass loss) in all regions except the two northernmost ones: kpc and thu. along the south-western ice sheet margin the 2015 ablation anomalies appear to be one third below average. however, what has to be accounted for in the interpretation is that the promice observational period contained several warm years and summers. figure 2 illustrates that there have been considerable fluctuations in atmospheric temperatures at greenland coastal sites with continuous records dating back to the 19th century. the promice observational period distinctly classifies as one with aboveaverage temperatures, both in the relatively warm south and colder north (fig. 2). this is also likely to imply above-average ablation, and thus biased promice ablation anomalies. it is therefore more insightful to evaluate recent ablation measurements in the context of a more representative placing greenland ice sheet ablation measurements in a multi-decadal context dirk van as, robert s. fausto, john cappelen, roderik s.w. van de wal, roger j. braithwaite, horst machguth and the promice project team* * charalampos charalampidis, jason e. box, anne m. solgaard, andreas p. ahlstrøm, konstanze haubner, michele citterio and signe b. andersen fig. 1. map of greenland with 2015 ablation anomaly values referenced to the 1961–1990 period at the lower (suffix l) promice weather station sites. black dots indicate dmi weather stations selected for this study. 250 km qas_l nuk_l kan_l thu_l upe_l kpc_l tas_l station nord tasiilaq pituffik upernavik kangerlussuaq nuuk qaqortoq 14% sco_l ittoqqortoormiit –13% 184% –18% –1% 56% 67% 20% 7272 climate. greenland studies often use the 1961–1990 period, during which the ice sheet is assumed to have been in nearsteady state (e.g. braithwaite et al. 1992; rignot et al. 2008). the main aim of this study is to reference promice-measured ablation to this 1961–1990 ‘climate normal’. present-day temperatures in a multidecadal perspective in order to determine the 1961–1990 reference climate, we need observational records spanning that period and recent years. the only continuous and (on these time scales) relevant greenland data series that exist have been recorded in coastal areas, by weather stations of the danish meteorological institute (dmi). for this study we selected those dmi time series that were gathered closest to the promice weather station sites and were initiated before 1961 (fig. 1). the earliest measurements (primarily of air temperature) were taken in the 1700s; several continuous records date back to the mid to late 1800s (fig. 2). ice sheet surface melting occurs predominantly in summer, so for the purpose of this study we calculated the average june–july–august (jja) temperature for each year, following e.g. braithwaite et al. (1992). temperature anomalies were calculated by subtracting the 1961–1990 jja average. we obtained monthly average data from the dmi technical report 15-04 (cappelen 2015) and supplemented these with 2015 data. table 1 shows that at all selected dmi sites the jja temperature during the promice observational period is higher than in the reference period. the smallest difference is found at qaqortoq in south greenland (0.8°c), and the largest at ittoqqortoormiit in the east (2.4°c). for the years with promice annual ablation values (2008–2015), only three out of a total of 64 station years (5%) had negative jja temperature anomalies indicative of conditions colder than during the reference period, emphasising the need for a well-defined context. measurements of ice ablation along the ice-sheet margin the first promice weather stations were established in 2007 (van as et al. 2011), thus providing annual net-ablation values since 2008 (i.e. end of melt season 2007 until end of melt season 2008). at any given site five to eight years’ worth of ablation data exist. in this study we only make use of the eight weather stations that are located closest to the ice-sheet margin (all with suffix ‘l’ for ‘lower’). at these elevations, summer ablation is much larger than winter accumulation, resulting in stronger correlations between net ablation and atmospheric temperature anomalies than higher on the ice sheet where ablation becomes an increasingly small contributor to the surface-mass budget. selecting the ‘lower’ promice stations provides us with a total of 56 ablation years. table 1. temperature and ablation statistics for the promice weather station sites temperature jja 1961–1990 (°c) 1.7 2.7 5.5 6.5 5.5 9.2 4.1 3.9 temperature jja 2008–2015 (°c) 2.8 5.0 6.7 7.3 7.5 10.5 6.3 5.3 annual net ablation promice (m ice eq.) 2.2 3.2 3.6 6.4 5.5 3.8 2.6 1.8 2015 ablation anomaly ref. to the promice average (%) 7 –1 –14 –20 –31 –34 –12 30 annual net ablation 1961–1990 (m ice eq.) 1.4 2.7 2.7 5.1 4.4 3.1 1.5 0.8 temperature sensitivity (line slope) (m ice eq. °c–1) 0.71 0.25 0.79 1.23 0.64 0.63 0.55 0.89 correlation of linear fit (r) 0.59 0.50 0.60 0.75 0.83 0.71 0.81 0.94 rmsd* of linear fit (m ice eq.) 0.5 0.3 0.6 0.9 0.6 0.5 0.3 0.2 uncertainty ablation calculation (m ice eq.) 0.6 0.4 0.7 0.9 0.6 0.6 0.5 0.4 2015 ablation anomaly referenced to 1961–1990 (%) 67 ± 40 20 ± 15 14 ± 25 –1 ± 18 –13 ± 15 –18 ± 19 56 ± 31 184 ± 43 kpc_l sco_l tas_l qas_l nuk_l kan_l upe_l thu_l * rmsd: root mean squared differences −12 −9 −6 −3 0 3 1860 1880 1900 1920 1940 1960 1980 2000 2020 te m p er at u re ( °c ) year upernavik nuuk qaqortoq promice period reference period fig. 2. annual (dots) and five-year (lines) running-mean temperatures at the three longest-running dmi measurement sites used in this study. arrows indicate the 1961–1990 reference period and the promice period (2007–present). 73 we supplement the promice data with older ablation observations gathered in close proximity to the current station locations, at identical elevations, and covering the entire ‘hydrological’ year, i.e. referenced to the end of the melt season. at four station sites, such measurements from before 2008 exist, namely at qas_l for 2001–2007 (e.g. podlech et al. 2004), at nuk_l for 1981–1987 (e.g. braithwaite et al. 1992), at kan_l since 1991 (van de wal et al. 2012) and at thu_l in 1954 (schytt 1955). these 40 historical measurements bring our grand total to 96 ablation years. present-day ablation in a multi-decadal perspective in this study we relate annual net-ablation values to jja temperatures following e.g. braithwaite et al. (1992), but without precipitation due to lacking dmi data. besides, differences in precipitation at the dmi and promice sites can be large due to spatial heterogeneity. in fig. 3 we plotted the annual net-ablation values against the temperature anomalies calculated from dmi weather stations in the region. at all sites, the ablation totals typically increase with temperature, as indicated by the linear least-squares fit lines. the slopes of these lines are the regional temperature sensitivities, which is relatively low at the sco_l site with 0.25 m ice equivalent (eq.) °c–1, and high for qas_l where roughly an additional 1.2 m of ice would ablate for every degree jja temperature increase (table 1). we find an average temperature sensitivity of the ice-marginal area of 0.71 ± 0.28 m ice eq. °c–1 (standard deviation given), rather similar to the value of 0.5 m water eq. °c–1 mentioned in e.g. braithwaite et al. (1992). a key element in this study is the ablation value at which the fitted lines intercept the 0°c temperature anomaly line, which for some sites requires extrapolation. the intercept values represent net ablation in the reference climate. most promice ablation values exceed the intercept values specific to their sites: ablation is larger in recent years than in the reference period. table 1 shows that the largest relative increase is found at thu_l, where we estimate annual net ablation to have increased by c. 120%. the reference period adjustment for promice ablation values is substantial at all sites (0.6–1.2 m ice eq., table 1). we need to be cautious in the interpretation of these results as the number of data points is still rather small. furthermore, one cannot expect a perfect correlation between ablation and temperature because: (1) melt is the result of a surplus in energy at the ice-sheet surface, of which only part is provided by atmospheric heat content, (2) the horizontal and vertical distances between the dmi and promice observation sites are considerable, and (3) winter accumulation is part of the net ablation signal but not a function of jja temperatures. as a measure of uncertainty of reference-period ablation due to the above, we calculate the root mean squared differences (rmsd) between the measured and calculated values. these range from 0.2 to 0.9 m ice eq. (table 1). to this we add a conservatively chosen measurement uncertainty in the annual net-ablation values of 0.3 m ice eq. (fausto et al. 2012), resulting in total uncertainties ranging from 0.4 to 0.9 m ice eq. the largest uncertainty is found for qas_l, likely due to interannual variability in winter accumulation and the station’s positioning in an irregular, crevassed terrain prior to its 1.5 km relocation in 2009. through the above re-referencing procedure the 2015, ablation anomalies become considerably larger (table 1, fig. 1). after reference adjustment, we find positive 0 2 4 6 8 −2 −1 0 1 2 3 4 a b la ti o n ( m i ce e q .) temperature anomaly (°c) kpc_l sco_l tas_l qas_l nuk_l kan_l upe_l thu_l fig. 3. measured annual net ablation from promice (dots) and other projects (circles) plotted against the regional temperature anomaly referenced to the 1961–1990 period. lines illustrate linear least-squares fits. 0 2 4 6 8 1860 1880 1900 1920 1940 1960 1980 2000 2020 a b la ti o n ( m i ce e q .) year kpc tas nuk upe sco qas kan thu fig. 4. estimated yearly (dots) and five-yearly (lines) net ablation for sites currently instrumented by promice. 7474 2015 anomalies for all sites except qas_l, nuk_l and kan_l, though these do not exceed their uncertainty ranges. the largest anomaly still occurs at thu_l, but is 184 ± 43% when referenced to the 1961–1990 climate, six times larger than the 30% value when referenced to the promice average. approximating past ablation the relations found between jja temperature and annual net ablation in fig. 3 can be used to estimate ablation from temperature in any year. in fig. 4 we used these functions for all summers for which dmi temperature data are available – though some of the earlier data were discarded to maintain continuity. we assume the uncertainties listed in table 1 to apply for these calculations as well, although it should be noted that the derived functions are less well constrained in the range with negative temperature anomalies (fig. 3). we conclude that at our study sites annual net ablation is likely to be larger in recent years than during any previous period in the instrumental era, covering up to 150 years. especially at the more northern locations we find that ablation increases in recent years are large. yet fig. 4 suggests that in southern greenland ablation peaked significantly around 1930. while most of greenland underwent relatively warm (summer) conditions in the 1930s (cappelen 2015), this was most notable at the more southern locations, resulting in amplified ablation values according to our estimates. jja temperatures were higher in 1928 and 1929 than in any other year of the qaqortoq record, both attaining values of 9.2°c. this suggests that ablation in those years may have exceeded the largest net ablation measured on the greenland ice sheet (9.3 m ice eq. at qas_l in 2010), although this is not beyond the uncertainty that accounts for important factors such as winter accumulation. acknowledgements this is a publication in the framework of the dancea-funded programme for monitoring of the greenland ice sheet (promice), in collaboration with several other projects. the kan stations are funded by the greenland analogue project (gap). stake measurements by utrecht university are funded through the nwo polar programme. references andersen, m.l. et al. 2015: basin-scale partitioning of greenland ice sheet mass balance components (2007–2011). earth and planetary science letters 409, 89–95. braithwaite, r.j., olesen o.b. & thomsen h.h. 1992: calculated variations of annual ice ablation at the margin of the greenland ice sheet, west greenland, 1961–1990. journal of glaciology 38 (129), 266–272. cappelen, j. (ed.) 2015: greenland – dmi historical climate data collection 1784–2014. danish meteorological institute technical report 15–04, 97 pp. dutton, a., carlson, a.e., long, a.j., milne, g.a., clark, p.u., deconto, r., horton, b.p., rahmstorf, s. & raymo, m.e. 2015: sealevel rise due to polar ice-sheet mass loss during past warm periods. science 349 (6244) aaa4019. enderlin, e.m., howat, i.m., jeong, s., noh, m.-j., van angelen, j.h. & van den broeke, m.r. 2014: an improved mass budget for the greenland ice sheet. geophysical research letters 41 (3), 866–872. fausto, r.s., van as, d., ahlstrøm, a.p. & citterio, m. 2012: assessing the accuracy of greenland ice sheet surface ablation measurements by pressure transducer. journal of glaciology 58 (212), 1144–1150. khan, s.a., aschwanden, a., bjørk, a.a.,wahr, j., kjeldsen, k.k. & kjær, k.h. 2015: greenland ice sheet mass balance: a review. reports on progress in physics 78, 046801. nghiem, s.v. et al. 2012: the extreme melt across the greenland ice sheet in 2012. geophysical research letters 39, l20502. podlech, s., mayer, c. & bøggild, c.e. 2004: glacier retreat, mass-balance and thinning: sermilik glacier, south greenland. geografiska annaler 86a, 305–317. rignot, e., box, j.e., burgess, e. & hanna e. 2008: mass balance of the greenland ice sheet from 1958 to 2007. geophysical research letters 35, l20502. schytt, v. 1955: glaciological investigations in the thule ramp area. report snow, ice, and permafrost research establishment 28, 88 pp. corps of engineers, u.s. army. tedesco, m. et al. in press: greenland ice sheet [in ‘state of the climate in 2015’]. bulletin of the american meteorological society 97(7). van as, d., fausto, r.s. & promice project team 2011: programme for monitoring of the greenland ice sheet (promice): first temperature and ablation records. geological survey of denmark and greenland bulletin 23, 73–76. van de wal, r.s.w., boot, w., smeets, c.j.p.p., snellen, h., van den broeke, m.r. & oerlemans, j. 2012: twenty-one years of mass balance observations along the k-transect, west greenland. earth systems science data 4, 31–35. authors’ addresses d.v.a., r.s.f., h.m., c.c., j.e.b., a.m.s., a.p.a., k.h., m.c. & s.b.a., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: dva@geus.dk h.m. also at department of geography, university of zurich, winterthurerstrasse 190, 8057 zürich, switzerland. k.h. also at natural history museum, copenhagen university, øster voldgade 5, dk-1350 copenhagen k, denmark. j.c., danish meteorological institute, lyngbyvej 100, dk-2100 copenhagen ø, denmark. r.s.w.v.d.w., institute for marine and atmospheric research, utrecht university, princetonplein 5, 3584 cc utrecht, the netherlands. r.j.b., school of environment, university of manchester, oxford road, manchester m13 9pl, uk. geological survey of denmark and greenland bulletin 11, 1-7 1 geological survey of denmark and greenland bulletin 11 · 2006 precambrian crustal evolution and cretaceous–palaeogene faulting in west greenland edited by adam a. garde and feiko kalsbeek geological survey of denmark and greenland danish ministry of the environment 2 geological survey of denmark and greenland bulletin 11 keywords archaean, cretaceous–palaeogene, deformation, faulting, geochronology, kangâmiut dykes, nagssugtoqidian, rinkian, west greenland cover flat-lying, grey archaean orthogneisses cut by palaeoproterozoic dolerite dykes (black), which have been deformed and rotated into nearparallellism with their host rocks during the nagssugtoqidian orogenesis. the whole succession is cut by lateto postkinematic nagssugtoqidian pegmatites (pink). south coast of island 11 km east of aasiaat in the northern nagssugtoqidian orogen. the nagssugtoqidian deformation of archaean rocks in this region is the main subject of mazur et al. (this volume) and van gool & piazolo (this volume). photo: adam a. garde. frontispiece: facing page archaean, synto postkinematic granitic sheets cutting grey orthogneiss and variably folded with their host. coastal exposure on the north coast of saqqarput 30 km south-east of kangaatsiaq, within an archaean block in the northern nagssugtoqidian orogen that has largely escaped palaeoproterozoic deformation. compare with the cover photograph displaying an example of palaeoproterozoic deformation, and see articles by mazur et al. and thrane & connelly (this volume). photo: adam a. garde. chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, geological institute, university of aarhus; minik rosing, geological museum, university of copenhagen; finn surlyk, geological institute, university of copenhagen scientific editors of this volume: adam a. garde and feiko kalsbeek editorial secretary: esben w. glendal referees: steffen bergh (norway), clark friend (uk), john grocott (uk), karen hanghøj (usa), paul martin holm (denmark), åke johansson (sweden), graham leslie (uk), kenneth j.w. mccaffrey (uk), allen p. nutman (australia), chris pulvertaft (denmark), andrew saunders (uk) and three anonymous referees illustrators: adam a. garde, eva melskens and henrik klinge-pedersen digital photographic work: benny m. schark graphic production: knud gr@phic consult, odense, denmark printers: schultz grafisk, albertslund, denmark receipt/acceptance dates of manuscripts: see end of individual articles printed: 5 december 2006 issn 1604-8156 isbn-10: 87-7871-188-6 isbn-13: 978-87-7871-188-5 geological survey of denmark and greenland bulletin the series geological survey of denmark and greenland bulletin replaces geology of denmark survey bulletin and geology of greenland survey bulletin. citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 11, 204 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk or geografforlaget aps rugårdsvej 55, dk-5000 odense c, denmark phone: +45 63 44 16 83, fax: +45 63 44 16 97, e-mail: go@geografforlaget.dk © danmarks og grønlands geologiske undersøgelse (geus), 2006 3 4 contents preface a.a. garde and f. kalsbeek . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5 evolution of neoarchaean supracrustal belts at the northern margin of the north atlantic craton, west greenland j.a. hollis, m. keiding, b.m. stensgaard, j.a.m. van gool and a.a. garde. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9 pre-nagssugtoqidian crustal evolution in west greenland: geology, geochemistry and deformation of supracrustal and granitic rocks north-east of kangaatsiaq j.-f. moyen and g.r. watt . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 33 a lead isotope study of an archaean gold prospect in the attu region, nagssugtoqidian orogen, west greenland h. stendal, r. frei and b.m. stensgaard . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 53 origin and evolution of the kangâmiut mafic dyke swarm, west greenland k.r. mayborn and c.e. lesher . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 61 zircon geochronology from the kangaatsiaq–qasigiannguit region, the northern part of the 1.9–1.8 ga nagssugtoqidian orogen, west greenland k. thrane and j.n. connelly . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 87 207pb-206pb dating of magnetite, monazite and allanite in the central and northern nagssugtoqidian orogen, west greenland h. stendal, k. secher and r. frei . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 101 new hornblende and muscovite 40ar/39ar cooling ages in the central rinkian fold belt, west greenland a.-s. sidgren, l. page and a.a. garde . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 115 presentation and interpretation of structural data from the nagssugtoqidian orogen using a gis platform: general trends and features j.a.m. van gool and s. piazolo . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 125 the nordre strømfjord shear zone and the arfersiorfik quartz diorite in arfersiorfik, the nagssugtoqidian orogen, west greenland k. sørensen, j.a. korstgård, w.e. glassley and b.m. stensgaard . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 145 structural analysis of the northern nagssugtoqidian orogen, west greenland: an example of complex tectonic patterns in reworked high-grade metamorphic terrains s. mazur, s. piazolo and g. i. alsop . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 163 magnetic anomalies and metamorphic boundaries in the southern nagssugtoqidian orogen, west greenland j.a. korstgård, b.m. stensgaard and t.m. rasmussen . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 179 faults and fractures in central west greenland: onshore expression of continental break-up and sea-floor spreading in the labrador – baffin bay sea r.w. wilson, k.e.s. klint, j.a.m. van gool, k.j.w. mccaffrey, r.e. holdsworth and j. a. chalmers . . . . . . . . . . 185 5 preface the present volume marks the completion of a large research project by the geological survey of denmark and greenland (geus), focused on the northern part of the palaeoproterozoic nagssugtoqidian orogen of central west greenland, and carried out by a team of danish and international participants. the project comprised geological mapping as well as structural, geochronological, geochemical and economic geological studies. this volume contains reports on both archaean and palaeoproterozoic geology as well as a study of neotectonic brittle structures. the field work was carried out in 2000–2003 in the region between nordre strømfjord and jakobshavn isfjord (see e.g. van gool & piazolo 2006, this volume, fig. 1). the project had two immediate purposes, namely to establish an overview of the mineral resource potential of supracrustal rocks in the region between 66° and 70°15′n, and produce four new geological sheets in the survey’s 1:100 000 map series. the first collection of papers about the nagssugtoqidian orogen, published by the geological survey of greenland (ggu, now part of geus), dates back to 1979 (korstgård 1979). the investigations in this period were mainly based on field descriptions and structural analysis of coastal areas in the southern and central parts of the orogen, combined with limited petrographical, palaeomagnetic and geochronological studies; the results also comprised the first 1:100 000 geological map from within the nagssugtoqidian orogen (olesen 1984). the proterozoic age of the orogen had been established, but it was believed that most, if not all of the quartzofeldspathic basement gneisses were of archaean origin. subsequent work in the nagssugtoqidian orogen by ggu in the 1980s showed that besides archaean orthogneisses and supracrustal rocks, the central part of the orogen also comprises the root zone of a palaeoproterozoic magmatic arc and associated panels of palaeoproterozoic volcanic and metasedimentary rocks (kalsbeek et al. 1987). these results were confirmed during further investigations by the danish lithosphere centre (dlc) in 1994–1999, and the plate-tectonic collisional history of the southern and central nagssugtoqidian orogen was described in detail (van gool et al. 2002). however, these studies added little to previous knowledge of the northern parts of the orogen in the kangaatsiaq–aasiaat–qasigiannguit region, knowledge that was largely based on coastal reconnaissance by henderson (1969) at the time when the entire orogen was still believed to consist of archaean rocks. another project preceding the present work was carried out by ggu in 1988–1991 immediately north of the nagssugtoqidian orogen, in the southernmost part of the likewise palaeoproterozoic rinkian fold belt (disko bugt project, kalsbeek 1999). it was shown that also the latter region comprises palaeoproterozoic (meta)sedimentary rocks, and that most of the archaean basement is strongly overprinted by palaeoproterozoic structures that were formed during overall wor nw-directed lateral tectonic transport. although these structures might be related to similar structures in the nagssugtoqidian orogen, the relationship between the nagssugtoqidian orogen and the rinkian fold belt remained speculative. the only previous economic geological study of regional extent in central west greenland was an airborne reconnaissance study supplemented by local field work, which was carried out in the early 1960s by kryolitselskabet øresund a/s. this work resulted in the discovery of a massive sulphide deposit at naternaq (lersletten), which was studied again in some detail in 2001 by the survey (østergaard et al. 2002) but not reported on in the present volume. the present volume comprises 12 papers with topics ranging geochronologically from mid-archaean to palaeogene, and geographically from the southern nagssugtoqidian foreland to the central part of the rinkian fold belt. many of the papers deal with the northern part of the nagssugtoqidian orogen and are related to the recent field work in that region, while a few contributions are rooted in dlcor other projects. the papers have been arranged in approximate chronological order and are grouped in terms of their main subjects. the two first papers, by hollis et al. and moyen & watt, deal with archaean supraand infracrustal rocks in the northern nagssugtoqidian orogen: their origin, ages, and structural and metamorphic evolution. these papers provide insight into the age and origin of the continental crustal orthogneisses and granites that underlie most of the region, and discuss the relationships between the supracrustal and plutonic components, using zircon u-pb age determinations and major and trace element geochemical characteristics. also the question of palaeoproterozoic tectonic overprint is discussed, with the conclusion © geus, 2006. geological survey of denmark and greenland bulletin 11, 5–7. available at: www.geus.dk/publications/bull 6 from both study areas that most of the observed structures are archaean. the third paper with focus on archaean geology, by stendal et al., describes a small gold prospect at attu likewise in the northern nagssugtoqidian orogen, and discusses the age of the prospect and its host rocks using pbpb geochronology of magnetite. it is concluded that the host rocks at attu may be as old as 3162 ± 43 ma, and that the gold prospect itself is around 2650 ma in age. the fourth paper, by mayborn & lesher, is a thorough review of the kangâmiut dyke swarm in the southern nagssugtoqidian orogen and its foreland. it includes new whole-rock and mineral chemical data, and a list of sampling sites and corresponding field data. the emplacement mechanism and depth of the dyke swarm are discussed in detail, and it is concluded that the dykes were emplaced during the initial rifting prior to the nagssugtoqidian collision and that they are unrelated to subduction processes (contrary to the belief by some previous authors). the next three papers provide geochronological constraints on the ages of supraand infracrustal rocks and the deformation and metamorphism in the northern nagssugtoqidian orogen, and on late orogenic uplift in the central rinkian fold belt. in the first of these papers thrane & connelly employ zircon u-pb age determinations (mainly using the laser icp-ms method), and for the first time provide unequivocal documentation that the naternaq supracrustal belt is of palaeoproterozoic age. other zircon age data from a synkinematic granite southeast of kangaatsiaq show that the large fold structures in this region are of archaean age. the subsequent paper by stendal et al. presents pb-pb ages and isotopic signatures of magnetite in amphibolites; the obtained ages are younger than 1800 ma and are related to cooling of the orogen. stepwise leaching pb-pb ages of monazite and allanite in pegmatites fall in the range of 1750–1800 ma, and are interpreted to date the emplacement of these rocks. the third paper in this group, by sidgren et al., deals with new 40ar/39ar ages of around 1790 ma (hornblende) and 1680 ma (muscovite) from archaean and palaeoproterozoic rocks in the central rinkian fold belt, which are interpreted as orogenic cooling ages. the hornblende ages are significantly older than such hornblende ages previously obtained from the central and northern nagssugtoqidian orogen, pointing to different uplift histories in the two regions. this may in turn suggest that the rinkian continental collision preceded that in the nagssugtoqidian orogen. four of the remaining five papers deal with the nagssugtoqidian structural evolution. in the first of these, van gool & piazolo present a new method of structural analysis, where a geographical information system (gis) is used as a framework for visualisation and analysis of large amounts of structural data. the paper graphically presents an overview of thousands of data points within an area of approximately 160 × 180 km in the central and northern parts of the nagssugtoqidian orogen. this interesting data set points directly towards the two next papers, where crustal-scale structures in the same region and their origin are discussed: sørensen et al. address the prominent nordre strømfjord shear zone just south of this block, and describes the structural and metasomatic transition into the shear zone by means of aeromagnetic and lithological map patterns and geochemical data. another paper, by mazur et al., addresses a prominent break in the structural pattern within the kangaatsiaq–aasiaat area, where the southern part acted as a rigid block during the nagssugtoqidian orogeny and thus preserved its archaean structure. the fourth paper in this group, by korstgård et al., combines rock and aeromagnetic data to discuss the relationship between structure, metamorphic facies and total magnetic field intensity anomalies in the southern nagssugtoqidian orogen. the authors show that static metamorphic boundaries are gradual, whereas boundaries along deformation zones are abrupt. the last paper, by wilson et al., is a novel remote sensing and field geological analysis of onshore brittle structures related to the complex ungava fault zone in the davis strait, which developed during the cretaceous–palaeogene opening of the labrador sea – davis strait – baffin bay seaway. the study area is located in the central nagssugtoqidian orogen, and the authors carefully establish a distinction between old nagssugtoqidian and younger structures in the basement rocks and identify five main sets of young lineaments. they conclude that the onshore fault patterns are predominantly of strike-slip nature, and that they reflect the stress fields that governed the opening of the seaway. acknowledgements the editors are grateful to the 14 external reviewers, each of whom reviewed one or more of the individual papers, for their thorough and constructive work. 7 references henderson, g. 1969: the precambrian rocks of the egedesminde–christianshåb area, west greenland. rapport grønlands geologiske undersøgelse 23, 37 pp. kalsbeek, f. (ed.) 1999: precambrian geology of the disko bugt region, west greenland. geology of greenland survey bulletin 181, 179 pp. kalsbeek, f., pidgeon, r.t. & taylor, p.n. 1987: nagssugtoqidian mobile belt of west greenland: a cryptic 1850 ma suture between two archaean continents – chemical and isotopic evidence. earth and planetary science letters 85, 365–385. korstgård, j.a. (ed.) 1979: nagssugtoqidian geology. rapport grønlands geologiske undersøgelse 89, 146 pp. olesen, n.ø. 1984: geological map of greenland, 1:100 000, agto 67 v.1 nord. copenhagen: geological survey of greenland. østergaard, c., garde, a.a., nygaard, j., blomsterberg, j., nielsen, b.m., stendal, h. & thomas, c.w. 2002: the precambrian supracrustal rocks in the naternaq (lersletten) and ikamiut areas, central west greenland. geology of greenland survey bulletin 191, 24–32. van gool, j.a.m. & piazolo, s. 2006: presentation and interpretation of structural data from the nagssugtoqidian orogen using a gis platform: general trends and features. in: garde, a.a. & kalsbeek, f. (eds): precambrian crustal evolution and cretaceous–palaeogene faulting in west greenland. geological survey of denmark and greenland bulletin 11, 125–144 (this volume). van gool, j.a.m., connelly, j.n., marker, m. & mengel, f.c. 2002: the nagssugtoqidian orogen of west greenland: tectonic evolution and regional correlation from a west greenland perspective. canadian journal of earth sciences 39, 665–686. adam a. garde and feiko kalsbeek geological survey of denmark and greenland bulletin 37, 1-74 + appendices geological survey of denmark and greenland bulletin 37· 2016 biostratigraphic correlation of the western and eastern margins of the labrador–baffin seaway and implications for the regional geology henrik nøhr-han sen, graham l. williams & robert a. fensome geological survey of denmark and greenland ministry of energy, utilities and climate bulletin37.qxp_bulletin 37 26/01/17 13.59 side 1 geological survey of denmark and greenland bulletin 37 keywords baffin margin, labrador margin, offshore west greenland, cretaceous, cenozoic, biostratigraphy, palynology, dinocysts, palaeoenvironments cover illustration the canadian icebreaker amundsen in blanley bay, southern devon island in the canadian arctic in the autumn of 2013. immediately south-east of this locality, across lancaster sound, is bylot island where cretaceous–paleocene sediments are exposed. photo: kate jarrett (gsc). frontispiece: facing page survey drilling team in agatdalen, central nuussuaq, west greenland, recovering core from the lower–middle campanian aaffarsuaq member (itilli formation); the mountain in the background is composed of paleocene volcanic rocks of the vaigat formation. such fully cored stratigraphic wells have been invaluable in documenting the detailed stratigraphy of the onshore area of the west greenland margin. chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, department of geoscience, aarhus university; minik rosing, geological museum, university of copenhagen; finn surlyk, department of geosciences and natural resource management, university of copenhagen scientific editor of this volume: jon r. ineson editorial secretary: jane holst referees: raquel guerstein (argentina) and gregers dam (denmark) illustrators: jette halskov (geus) and bill macmillan (gsca) digital photographic work: benny m. schark graphic production: annabeth andersen printers: rosendahls · schultz grafisk · denmark manuscript received: 21 january 2015 final version approved: 17 march 2016 printed: 28 december 2016 issn (print) 1604-8156 issn (online) 1904-4666 isbn (print) 978-87-7871-445-9 isbn (online) 978-87-7871-447-3 citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bulletin 37, 74 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk or at www.geus.dk/publications/bull © de nationale geologiske undersøgelser for danmark og grønland (geus), 2016 for the full text of the geus copyright clause, please refer to www.geus.dk/publications/bull bulletin37.qxp_bulletin 37 26/01/17 13.59 side 2 bulletin37.qxp_bulletin 37 26/01/17 13.59 side 3 4 contents abstract . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . tectonic setting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . labrador sea . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . davis strait and baffin bay . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . general stratigraphy . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . offshore . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . labrador margin . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . west greenland margin . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . baffin margin . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . onshore . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . west greenland . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . north-eastern canada . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . methodology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . previous palaeontological studies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . biostratigraphic results . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . lower cretaceous . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . aptian . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . albian . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . upper cretaceous . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . cenomanian . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . turonian–santonian. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . campanian . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . maastrichtian . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . palaeogene . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . danian–selandian . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . thanetian . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . ypresian . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . lutetian . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . bartonian . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . priabonian . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . rupelian . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . chattian . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . neogene . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . miocene . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . pliocene–pleistocene . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . cretaceous . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . palaeogene . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . neogene . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . palaeoenvironmental results . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . general considerations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . palaeoenvironments and dinocysts: previous studies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . previous palaeoenvironmental interpretations from the labrador–baffin seaway . . . . . . . results and interpretations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . cretaceous . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . cenozoic . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7 9 11 11 11 12 12 12 13 16 16 16 17 18 20 21 21 21 23 23 23 23 24 25 25 25 26 28 29 29 29 30 30 30 30 31 32 32 32 33 36 36 37 39 39 39 40 bulletin37.qxp_bulletin 37 26/01/17 13.59 side 4 palaeogeography, palaeoclimatology and palaeoceanography . . . . . . . . . . . . . . . . . . . . . . conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . references . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . appendix 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . appendix 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . appendix 3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5 44 48 49 50 60 67 74 bulletin37.qxp_bulletin 37 26/01/17 13.59 side 5 bulletin37.qxp_bulletin 37 26/01/17 13.59 side 6 7 abstract authors’ addresses h.n.-h., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: hnh@geus.dk g.l.w. & r.a.f., geological survey of canada (atlantic), natural resources canada, po box 1006, 1 challenger drive, dartmouth, nova scotia b2y 4a2, canada. new analyses of the palynological assemblages in 13 offshore wells on the canadian margin and six on the west greenland margin, in conjunction with onshore data, have led to a new biostratigraphic framework for the cretaceous–cenozoic strata of the labrador sea – davis strait – baffin bay (labrador–baffin seaway) region and the first broad biostratigraphic correlation of the canadian and greenland margins. this framework is based on 167 last occurrences and 18 local/regional peak/ common-occurrence events for dinocysts, miospores, fungal spores and azolla. detailed biostratigraphic evidence has confirmed the following hiatuses: pre-aptian in the hopedale basin; pre-albian in the saglek basin; albian–turonian in some wells of the hopedale basin; turonian–santonian/campanian in some areas; pre-campanian and late campanian – thanetian on the greenland margin; late maastrichtian and danian in some wells of the hopedale basin and in the saglek basin; selandian in part of the hopedale basin, in all the saglek basin wells and in two wells on the west greenland margin; late ypresian and/or lutetian on both sides; oligocene to middle miocene of considerable variability on both margins, with all of the oligocene and the lower miocene missing in all the west greenland margin wells; and middle to late miocene on the western side. on the canadian margin, the hiatuses can be partially matched with the five previously recognised regional unconformities; on the greenland margin, however, the relationship to the five unconformities is more tenuous. palyno morph assemblages show that most aptian to albian sediments were deposited in generally non-marine to marginal marine settings, interrupted by a short-lived shallow marine episode in the aptian. a marine transgression started in the cenomanian–turonian and led to the most open-marine, oceanic conditions in the campanian–lutetian; shallowing probably started in the late lutetian and continued into the rupelian, when inner neritic and marginal marine palaeoenvironments predominated. throughout the rest of the cenozoic, inner neritic palaeoenvironments alternated with marginal marine conditions on the margins of the labrador–baffin seaway. these observations broadly reflect the tectonic evolution of the seaway, with rift conditions prevailing from aptian to danian times, followed by drift through much of the paleocene and eocene, and post-drift from oligocene to the present. dinocysts indicate that climatic conditions in the labrador–baffin seaway region were relatively temperate in the cretaceous, but varied dramatically through the cenozoic. the danian was a time of increasingly warmer climate, a thermal maximum being reached around the paleocene–eocene boundary reflecting the global thermal event at this time. warm to hot conditions prevailed throughout the ypresian, but the climate began to cool in the lutetian, a trend that accelerated through the priabonian and rupelian. throughout the neogene, temperatures generally declined, culminating in the quaternary. nøhr-hansen, h., williams, g.l. & fensome, r.a. 2016: biostratigraphic correlation of the western and eastern margins of the labrador–baffin seaway and implications for the regional geology. geological survey of denmark and greenland bulletin 37, 74 pp. bulletin37.qxp_bulletin 37 26/01/17 13.59 side 7 8 r canadacanada ellesmere island 70°n 60°n 45°w55°w65°w labrado r nuussuaq nuussuaq basin hareøen disko bylot island cape dyer scott inlet buchan gulf quqaluit/ padloping islands saglek b asin 646 112 113 645 umiivik-1 delta-1 t4-1 t8-1 alpha-1/s1 gamma-1 hellefisk-1 at7-1 at2-1 lf7-1 ikermiut-1 nukik-2 nukik-1 qulleq-1 gjoa g-37 ralegh n-18 hekja o-71 rut h-11 gilbert f-53 karlsefni a-13 skolp e-07 pothurst p-19 ogmund e-72 snorri j-90 bjarni o-82 bjarni h-81 herjolf m-92 north bjarni f-06 roberval k-92 north leif i-05 cartier d-70 hare bay e-21 freydis b-87 south labrador n-79 kangâmiut-1 647 hopedalebasin cum berland sound h om e bay b af f in is lan d canada d a v is s t ra i t la b ra d o r s e a gro#3 250 km baffin bay greenland h udson stra i t 75°w shallow cored borehole greenland–canada border odp or dsdp borehole exploration well with gas shows exploration well exploration well with oil and gas shows 500 km greenland labrador sea baffin bay d avis strait nar es st ra it b a ffin isla n d 60°n 60°w 40°w 70°n 80°n bulletin37.qxp_bulletin 37 26/01/17 13.59 side 8 canada and greenland are separated, from south to north, by the labrador sea, the davis strait, baffin bay (referred to here collectively as the labrador–baffin sea way), and the narrow nares strait (fig. 1). as sum marised in monger et al. (2014) and fensome et al. (2014), the labrador–baffin seaway originated as a series of rift basins that developed successively from south to north during the early cretaceous, and which ultimately became connected to form a seaway in the late cretaceous. a drift phase in the seaway, causing the greenland plate to rotate away from north america, was heralded by an episode of volcanic activity between 62 and 56 million years ago, perhaps resulting from the passage of the icelandic mantle plume (dam et al. 1998a; larsen et al. 2016). this episode involved vast outpourings of basalt in central west greenland and south-eastern baffin island. as drift ensued, sea-floor spreading resulted in progressive widening of the la b rador–baffin seaway until the priabonian. this development meant that greenland changed course, its northern end colliding with arctic canada, resulting in an episode of deformation known as the eurekan oro g eny. the collision inhibited the rotation of green land away from north america, and sea-floor spreading in the labrador–baffin seaway ceased around the latest eocene. from that time on, sea-floor spreading in the northern north atlantic ocean occurred only between greenland and northwestern europe, where a spreading axis had opened about 55 million years ago. thus, until this time, greenland was part of laurasia, be tween about 55 and 35 million years ago it constituted a separate plate, and after 35 million years ago it became part of the north american plate. the labrador–baffin seaway stretches from about 52°n to 75°n, a distance of roughly 2500 km (fig. 1). knowledge of the timing of the geological evolution of the seaway is based primarily on biostratigraphic anal yses on exploration wells drilled between 1971 and 2000 on the labrador margin, offshore west green land and in the davis strait, and from some shallow cored boreholes drilled in baffin bay in the 1970s and 1980s and in the nuus suaq basin in the 1990s (fig. 1). additional information comes from numerous creta ceous–ceno zoic out crop sections in the nuussuaq basin in west green land, a few onshore sections in labrador and on baffin and bylot islands, and from odp leg 105, hole 645 in baffin bay (fig. 1). although previous work on these wells and sections has provided a good stratigraphic foundation for the cretaceous–cenozoic of the region, renewed petroleum exploration interest in recent years has revealed gaps in our knowledge and a need to correlate the western and eastern margins of the seaway. consequently, for the past decade the geological survey of denmark and greenland (geus) and the geo logical survey of canada (atlantic) (gsca) have undertaken an exhaustive palynological study, with the goal of providing more detailed age control. this study is based on palynological analysis of more than 2000 well samples from the labrador margin, the davis strait and offshore west greenland. organic-walled dino flagellate cysts (dinocysts) are the primary palyno morph group eva l ua ted, but spores and pollen (mio spores), a fern micro spore massula (azolla), algal and fungal microfossils and acritarchs have also been considered. the new age determinations presented here and documented in detail in a companion publication (fensome et al. 2016) are more precise than those of previous studies because of advances made in refining the stratigraphic ranges of dinocysts (e.g. williams et al. 2004). moreover, the number of species for which detailed stratigraphic information is available has increased immeasurably since early studies in the region in the 1970s. using the more detailed stratigraphic data, 187 bioevents are identified for the labrador–baffin seaway. some events are coeval, so the 187 bio events define 106 bioevent horizons. bioevent data are also incorporated from piasecki (2003) for the neo gene of offshore west greenland and from pedersen & nøhr-hansen (2014) for albian to paleocene strata of the nuussaq basin, west greenland. in addition to biostratigraphic determinations, palaeoenvironmental inferences are made from the data, using, for example, individual dinocyst species and dinocyst assemblages. we have also endeavoured to decipher palaeo oceanographic conditions and palaeoclimates, the results revealing similarities to high southern latitudes in the late cretaceous. 9 introduction facing page: fig. 1. map of the labrador–baffin seaway showing the location of relevant wells, boreholes and onshore localities; wells and boreholes shown in red were used in this study. inset map shows the regional context of the labrador–baffin seaway between canada and greenland. bulletin37.qxp_bulletin 37 26/01/17 13.59 side 9 10 rre s es s t ai t island ■■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■■ ■■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ▲ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■ ■ ■ ■ ■■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■ ■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■■ ■■ ■■ ■ ■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■■ ■ ■■ ■ ■■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■■ 646 112 113 645 rut h-11 gilbert f-53 karlsefni a-13 pothurst p-19 skolp e-07 ogmund e-72 snorri j-90 bjarni o-82 roberval k-92 north leif i-05 freydis b-87 cartier d-70 herjolf m-92 north bjarni f-06 250 km south labrador n-79 647 saglek basin hellefisk-1 ikermiut-1 nukik-2 delta-1 t4-1 t8-1 gamma-1 alpha-1/s1 nukik-1 qulleq-1 lf7-1 at2-1 at7-1 gjoa f-37 ralegh n-18 hekja o-71 kangâmiut-1 nuussuaq svartenhuk halvø disko cape dyer umiivik-1 gro#3 45°w 35°w sediments, large basins (400–0 ma) oceanic crust, eocene oceanic crust, age uncertain basalts and intrusives (60–30 ma) basement locally covered by sediment spreading axis extensional fault compressional fault thrust transition fault odp or dsdp borehole oceanic crust, paleocene exploration well exploration well with gas shows, exploration well with oil and gas shows shallow cored borehole fault (undifferentiated) paam iut b asin n u u k b as in k an gâ m iu t b as in lady franklin basin sisimiut basin disko bugt ik er m iu t b as in h o m e b ay b asin lancaster basin eclipse trough m elville bay g raben buchan g raben scott g raben d av is s tr ai t h ig h baffin basin n u u ss u au q b as in labrado r labrador sea hudson strait c um berland sound m elville bay k ivioq basin canada greenland bylot island baffin island 65°w 55°w 45°w 60°n 70° 70°n 55°w65°w75°w h opedale basin hawke basin west greenland volcanic province bulletin37.qxp_bulletin 37 26/01/17 13.59 side 10 11 labrador sea kerr (1967) and grant (1980) considered the labrador sea to be floored by foundered continental crust. umpleby (1979) explained the formation of this sea way by using a modification of the undation theory of van bemmelen (1949, 1972), but still considered that there was no underlying oceanic crust. the seismic, mag netic and gravity data (le pichon et al. 1971; srivastava 1978), however, strongly supported a platetectonic model of lithospheric extension and spreading, accompanied by generation of oceanic crust. srivastava (1978) considered the oldest oceanic crust in the southern labrador sea to be assignable to anomaly 34 (coniacian–santonian) and in the north to anomaly 31 (maastrichtian). he observed that a change in the direction of spreading occurred between anomaly 25 (early thanetian) and anomaly 24 (late thanetian to early ypresian), and noted that anomaly 13 (late pria bonian to earliest rupelian) and subsequent anomalies were missing. roest & srivastava (1989) and srivastava & roest (1999) interpreted the oldest oceanic crust in the la b rador sea as anomaly 33 (most of the campanian; 73.6–79 ma). however, biostratigraphic analyses of the qulleq-1 well, offshore west greenland, demonstrated that the well reached total depth (td) in santonian sedimentary rocks (nøhr-hansen et al. 2000; christi ansen et al. 2001; henriksen et al. 2009), which would imply that the crust underlying qulleq-1 is the oldestknown oceanic crust north of 56°n in the atlantic (henriksen et al. 2009). however, chalmers (1991) and chalmers & laursen (1995) could not identify any anomalies in the labrador sea older than chron 27n (latest danian) and concluded that sea-floor spreading began between 63 and 61.7 ma (latest danian), and that a transition zone exists between continental and oceanic crust. the interpretations of chalmers (1991) were later supported by those of chalmers & pulvertaft (2001) and oakey & chalmers (2012), and followed by sønder holm et al. (2003a, b). moreover, funk et al. (2007, 2012) suggested that the crust underlying qulleq-1 is of continental origin. based in part on the understanding of magnetic ano malies at the time and in part on biostratigraphy, balk will & macmillan (1990) proposed three megasequences for the lithostratigraphic units encountered in offshore wells: synrift (neocomian to campanian), drift (late cre taceous to eocene), and post-drift (oligocene to recent). similarly, but based on updated knowledge of the timing of sea-floor spreading, sønder holm et al. (2003b) recognised: an initial rifting phase (phase 1) in the early cretaceous, a continued rifting phase (phase 2) in the late cretaceous to danian, and a subsequent drift and post-drift phase (phase 3). sønderholm et al. (2003b) speculated that phase 1 may have started in the jurassic based on tentative observations by dalhoff et al. (2006) and piasecki (2003), a notion supported by larsen et al. (2009) who indicated that lithospheric stret ching between canada and west greenland probably started in jurassic time. on the western margin of the labrador sea from south to north is a series of roughly se–nw-aligned basins separated by basement highs, or arches, oriented perpendicular to the basins (fig. 2). exploration wells have been drilled in two of the basins, the hopedale and saglek basins. davis strait and baffin bay the davis strait high, where water depths can be as shallow as about 600 m, forms the northern periphery of the labrador sea and sits above a major transform margin that developed around the paleocene–eocene transition (oakey & chalmers 2012). today, the davis strait connects the labrador sea with baffin bay. as was the case for the labrador sea, both non-plate-tec tonic and plate-tectonic models have been proposed for baffin bay. kerr (1967) and grant (1975) considered baffin bay to be underlain by continental crust modified by basalt intrusions. keen et al. (1972) and sriva stava (1978), however, postulated that oceanic crust tectonic setting facing page: fig. 2. map of the basic geology of the labrador–baffin seaway, showing the main structural features and relevant well/borehole locations; wells and boreholes shown in red were used in this study. modified from henriksen et al. (2009) and oakey & chalmers (2012). bulletin37.qxp_bulletin 37 26/01/17 13.59 side 11 12 general stratigraphy underlies baffin bay. support for a plate-tectonic origin was provided by jackson et al. (1979), who identified anomalies 21 (earliest lutetian) to 13 (late priabonian to earliest rupelian). oakey (2005) recognised two sets of linear magnetic anomalies trending nnw–sse and nw–se, which correspond respectively to the paleo cene and eocene anomalies in the labrador sea. har rison et al. (2011) considered baffin bay to be under lain in part by oceanic crust, formed between chron 27 and chron 13 – i.e. late danian to earliest rupelian (fig. 2). in the east of baffin bay, off west greenland, is a series of roughly se–nwand s–n-aligned basins separated by basement highs or ridges oriented parallel or oblique to the basins (fig. 2). some of these basins along the south-western part of the west greenland margin were penetrated by wells drilled between 1976 and 2000: qulleq-1 in the paamiut basin, nukik-1 and nukik-2 in the nuuk basin, kangâmiut-1 in the kan gâ miut basin and ikermiut-1 in the sisimiut basin. another well, hellefisk-1, was drilled in the same period within the west greenland volcanic province (figs 1, 2). the oil company cairn drilled eight exploration wells in 2010 and 2011: one in the lady franklin basin (lf7-1), two in the nuuk basin (at2-1, at7-1), and five in the west greenland volcanic province (delta-1, t4-1, t8-1, alpha-1/s1, gamma-1; fig. 2); none of the samples or data from these wells are yet released. offshore labrador margin a formal cretaceous–cenozoic lithostratigraphic scheme was proposed for labrador margin wells by umpleby (1979) and modified by mcwhae et al. (1980); the ver sion adopted here (fig. 3) is that with the timescale calibrated by sørensen (2006, fig. 2) and modified by dickie et al. (2011). the oldest rocks in some wells in the hopedale and saglek basins are igneous or metamorphic and of pre cambrian age. in other wells in the two basins, the oldest rocks are ordovician or carboniferous. separa ting these basement rocks from overlying cretaceous and cenozoic strata is an unconformity, termed the labrador uncon formity by mcwhae et al. (1980) and mcwhae (1981; fig. 3). mcwhae (1981) gave an age of 130 to 120 ma for the labrador unconformity though the timescale on which this was based is not clear. the oldest cretaceous unit is the alexis for mation, which consists of basalts with subordinate claystone, siltstone and sandstone (fig. 3). corgnet & mcwhae (1973) determined radiometric ages of 139 ± 7 ma and 122 ± 6 ma for two basaltic cores from the type section of the alexis for mation in bjarni h-81, viz. valanginian to early aptian on the timescale of gradstein et al. (2012). sørensen (2006, fig. 2) gave an age of barremian to al bian for the bjarni formation, which overlies the alexis formation. the bjarni formation is a predominantly sand stone unit and the main hydrocarbon reservoir in the hopedale basin (figs 2, 3). according to mcwhae (1981), the bjarni formation is separated from the overlying markland formation by the avalon unconformity, which mcwhae considered to range from 100 to 85 ma (timescale not specified). the markland formation con sists mainly of a distal mudstone facies and its proximal sandstone equivalent, the freydis member. sønderholm et al. (2003b) determined an age of cenomanian to earliest paleocene for the markland formation. a sandstone unit at the top of the markland formation, the informal lower gudrid member, straddles the cretaceous–palaeo gene boundary (sørensen 2006). overlying the markland formation is the cart wright formation, which is predominantly a mudstone unit of paleocene age (sørensen 2006). in the saglek basin, the cartwright formation has lower and upper proximal sandstone facies, informally termed the mid dle and up per gudrid members respectively (balkwill & mcmillan 1990). separating the middle and upper gudrid members is the bylot unconformity; the duration of the hiatus was dated by mcwhae (1981) as being danian, 60–63 ma (timescale not specified) and which dickie et al. (2011) placed at the danian–selan dian boundary. bulletin37.qxp_bulletin 37 26/01/17 13.59 side 12 13 a major volcanic episode occurred in the northern labrador sea, the davis strait and southern baffin bay (figs 2, 3) during the paleocene to early eocene. in the rut h-11 well in the saglek basin, supposed tuff beds from this volcanic episode are reported to underlie the upper cretaceous – danian markland formation, yet have been dated as 59 ma (klose et al. 1982). a thick sequence of basalts in gjoa g-37 appears to be coeval with the tuff beds in rut h-11, with two basalt samples dated at 59.5 ± 1.0 ma and 59.2 ± 1.8 ma (williamson et al. 2001). in gjoa g-37, however, the basalts occur above the markland formation; the implication is that the volcanic rocks in rut h-11 are of intrusive origin. also of closely similar age are the cape dyer volcanics, dated at 58 ± 2 ma (clarke & upton 1971). william son et al. (2001) gave ages of 62.9 ± 2.5 ma and 55.1 ± 2.3 for two basalt cores recovered off cape dyer. these volcanics may be related to the transform margin that was developing in the northern saglek area through the davis strait, or to the onset of sea-floor spreading in southern baffin bay. on the labrador margin, the top of the informal upper gudrid member, and thus of the cartwright for mation, roughly corresponds to the paleocene–eocene boundary. successively above these rocks are the ken amu (eocene), mokami (oligocene to miocene) and sag lek (middle miocene to pleistocene) formations. each of these units consists broadly of a mudstone se quence that becomes sandier and even conglomeratic upwards. mcwhae et al. (1980) recognised a regional unconformity, the baffin bay unconformity, between the kenamu and mokami for mations. these authors also named the regional un con formity between the mo kami and saglek formations as the beaufort uncon formity. mcwhae (1981) considered the hiatus at the baffin bay unconformity to be within the early oligo cene, 38–34 ma; he placed the beaufort unconformity in the early to middle miocene, between 20 and 15 ma (timescale not specified). west greenland margin rolle (1985) formally defined seven formations from well sections released at that time in offshore west greenland (fig. 3). he assigned the oldest rocks, a sequence of interbedded mudstones and sandstones, to the narssarmiut formation and considered this to be equivalent to the freydis member of the markland formation to the west, and thus of campanian age. based on this study, the narssarmiut formation is now considered to be of selandian age (fig. 3). according to rolle (1985), the narssarmiut for mation is unconformably overlain by the ikermiut for mation, which comprises carbonaceous mudstone with some siltstone and sandstone. rolle (1985) considered the ikermiut formation to be of campanian to middle eocene age and correlated it with the markland and cartwright formations of the labrador margin. if these correlations are correct, then the ikermiut for mation would straddle the bylot unconformity. in deed, sønderholm et al. (2003b) illustrated the pre sence of a major hiatus within the ikermiut formation, possibly spanning the late campanian to earliest paleocene and thus equivalent to the bylot unconformity. the hellefisk formation occurs only in the hellefisk-1 well. rolle (1985) dated it as thanetian to eocene and considered that it may interfinger laterally with the ikermiut formation; we consider it to be a correlative of the upper part of the ikermiut formation (fig. 3). palaeogene formations of the west greenland mar gin are all siliciclastic, with mudstones dominant. the ataneq formation is mudstone-dominated; sandstonedominated units are the nukik, kangâmiut and manît soq formations. current age determinations of all these formations in individual wells are shown in fig. 3, based on work by dalhoff et al. (2003), nøhr-hansen (2003), rasmussen et al. (2003), rasmussen & sheldon (2003), sheldon (2003) and the results of this study. neogene sediments are generally upward-coarsening sequences of marine mudstones interbedded with wellsorted fine-grained sandstones. although the lithostratigraphy of the west green land margin confirmed in well sections extends back only to the santonian, older mesozoic strata are evident from seismic data. three seismic sequences have been recognised within the cretaceous succession (fig. 3). from oldest to youngest, these are the kitsissut (which probably includes volcanics), appat and kangeq se quences (chalmers et al. 1993, 1995; chalmers & pul vertaft 2001). the kitsissut sequence may be coeval with the lower part of the bjarni formation, with the inferred volcanics equivalent to the alexis formation. the kangeq sequence has been partly drilled in the ikermiut-1 and qulleq-1 wells, where the succession en countered is of early cam panian (sønderholm et al. 2003b) and san to nian–cam panian (christiansen et al. 2001) age, re spec tively. the mudstone-dominated kang eq sequence is the lithological equivalent of the markland formation of the lab rador margin. bulletin37.qxp_bulletin 37 26/01/17 13.59 side 13 14 upper cretaceous paleocene eocene oligocene miocene pliocene quaternary lower cretaceous 0 ma 5 10 15 20 25 30 35 40 45 50 55 60 65 70 75 80 85 90 95 100 105 110 115 120 125 130 135 140 145 aptian barremian hauterivian valanginian berriasian albian cenomanian turonian coniacian santonian campanian maastrichtian danian selandian thanetian ypresian lutetian bartonian priabonian rupelian chattian aquitanian burdigalian langhian serravallian tortonian messinian zanclean piacenzian c28 c29 c30 c31 c32 c33 c34n c34n c34n c34n m"-3"r m"-2"r m"-1"r c27 c26 c25 c24 c22 c21 c20 c19 c18 c17 c23 c16 c15 c13 magnetostrat. (polarity, chron) chronostratigraphy unconformitiesshelf basin labrador marginsw ne labrador trough erosion leif mb upper freydis mb lower freydis mb upper and middle gudrid mb lower gudrid mb sa gle k f m mokami fm lower kenamu fm cartwright fm lower lower upper upper upper bjarni fm alexis fm markland fm (labrador u nc.) (avalon unc.) (bylot unc.) bylot avalon labrador baffin bay beaufort (baffin bay unc.) (b eaufort u n c. ) fig. 3. stratigraphic framework of the mesozoic–cenozoic rocks of the labrador and west green land margins and adjacent on shore sections; modified from gregersen et al. (2013). the la brador margin stratigraphy is from dickie et al. (2011), with unconformities from mcwhae (1981). the west greenland mar gin stratigraphy is based on well stratigraphic studies by rolle (1985), nøhr-hansen (2003) and sønderholm et al. (2003); the deeper sub-well section is based on chalmers et al. (1993), chal mers & pulvertaft (2001) and sørensen (2006). the nuussuaq basin stratigraphy is from storey et al. (1998), dam et al. (2009), pedersen & nøhr-hansen (2014) and larsen et al. (2015); these pa pers and this study are the source of the inferred ages. the southeast baffin island stratigraphy is from burden & langille (1990) and pedersen et al. (2002), the home bay and scott inlet seabed samples from maclean et al. (2014) and the north baffin is land stratigraphy is based on jack son et al. (1978) and mcwhae (1981). the bylot island strati graphy is from miall (1986), waterfield (1989), and harrison et al. (1999). the timescale (ma) and magnetostratigraphy are from gradstein et al. (2012). fm: for mation (formal). fm: formation (informal). mb: member (for mal). mb: member (informal). unc.: unconformity. bulletin37.qxp_bulletin 37 26/01/17 13.59 side 14 15 h el le fis k1 g r o # 3 u m iiv ik -1 ik er m iu t1 k an gâ m iu t1 n u ki k2 n u ki k1 q u lle q -1 west greenland margin kangilia fm nuussuaq basin se baffin islandnwses n bylot island itilli fm atane fm ? itilli fm (ikermiut fm) kangeq sequence ikermiut fm ikermiut fm narssarmiut fm kangâmiut fm kangâmiut fm kangâmiut fm kangâmiut fm hellefisk fmnukik fm nukik fm ? ? ? ? ? ? ? ? ? ? ? ataneq fm ataneq fm manîtsoq fm manîtsoq fm manîtsoq fm fylla sand erqua fm naqerlog fm hareøen fm (hareøen) youngest volcanics quqaluit fm cape searle fm n u u ss u aq g ro u p (home bay – scott inlet) cape dyer basalt maligât fm svartenhuk fm vaigat fm atanikerluk fm eqalulik fm quikavsak fm kome fm basalts? sills/dykes appat sequence kitsissut sequence ? ? upernivik næs fm slibestensfjeldet fm hassel fm (labrador unc.) sermilik fm kanguk fm aktineq fm pond inlet fm navy board fm bylot island fm agatdal fm ikermiut fm mainly marine mudstones, locally sandy/silty mainly marine or deltaic sandy/silty deposits, locally mudstones potential and known source rocks volcanics td in basement unconformity mainly continental deposits bulletin37.qxp_bulletin 37 26/01/17 13.59 side 15 16 baffin margin offshore rocks of western baffin bay are known only from shallow cored boreholes drilled on the baffin shelf. according to maclean et al. (2014), cretaceous rocks of aptian–albian to cenomanian age occur on the south-eastern baffin shelf, and of coniacian to campanian age on the north-eastern baffin shelf. the former equate to the bjarni formation, the latter to the markland formation. however, according to maclean et al. (2014), danian to selandian rocks, of similar age and lithology to the eureka sound group (a unit generally used for strata farther north and west, see below) have been recovered from cores off cumberland sound (maclean & williams 1983; maclean et al. 1986). onshore west greenland mesozoic–cenozoic rocks in west greenland occur in the nussuaq basin, adjacent to baffin bay, where rocks of the nuussuaq group outcrop on several peninsulas and islands. dam et al. (2009) provided a synthesis of this group, in which they formalised the lithological units proposed by previous researchers, especially hen derson et al. (1976), and proposed some new units. stra tigraphic control in dam et al. (2009) was based on palynological studies, including those of koppelhus & pedersen (1993), nøhr-hansen (1993b, 1996), nøhrhansen & dam (1997), dam et al. (1998b, c), lans torp (1999), kennedy et al. (1999), nøhr-hansen & heilmann-clausen (2000), nøhr-hansen et al. (2002), sønderholm et al. (2003a, b). additional data were presented by pedersen et al. (2013) and pedersen & nøhr-hansen (2014). in the nuussuaq basin, the oldest known rift to early post-rift sediments are aptian? to albian in age and consist of fan-delta, fluvio-deltaic and shallow marine deposits of the kome formation, which onlap pre cam brian basement on disko and nuussuaq (hen derson et al. 1976; dam et al. 2009; pedersen & nøhr-hansen 2014). the succeeding slibestensfjeldet for mation com prises lacustrine or lagoonal deposits, with palynofloras dominated by non-marine miospores and rare, brackish-water dinocysts of albian, possibly mid dle albian, age (pedersen & nøhr-hansen 2014). both the sli be stensfjeldet and the kome formations are un con form ably overlain by albian to upper santonian post-rift, fluvial and wave-dominated deltaic to fully marine deposits of the atane formation (pedersen & pulver taft 1992; nøhr-hansen et al. 2002; dam et al. 2009; pedersen & nøhr-hansen 2014). to the west of a major fault, sediments of the atane formation are replaced by deep-water submarine-fan deposits as signed to the itilli formation. turonian – lower maa strichtian slope deposits are also described from svartenhuk halvø (dam et al. 1998b). lithological and age similarities suggest that the kome, slibestensfjeldet and bjarni formations are coeval, the unconformity beneath the atane formation being equivalent to the avalon unconformity. a major angular unconformity separates the deltaic deposits of the qilakitsoq member (the upper part of the atane formation) from the marine gravity-flow deposits of the lower–middle cam panian aaffarsuaq member of the itilli formation in nuussuaq (dam et al. 2000, 2009). rifting continued through the maastrichtian and into the danian in the nuussuaq basin, with at least three tectonic phases associated with valley and sub marine canyon incision (dam et al. 1998a; dam et al. 2009). the fill of the incised valleys represents three formations that are overlain by deep-water marine mudstones and volcanic tuffs (dam et al. 2009). al though occurrences of neogene rocks in west green land are sparse, a thin succession of neogene sediments overlies the eocene hareøen formation (fig. 3; hald 1976; christiansen et al. 1999). an insightful study of the rocks from onshore west greenland was undertaken by larsen (2006) and lar sen et al. (2009), who mapped and dated the dyke swarms in that region. these authors recognised the following phases of intrusion, which can be related to significant stages in the development of the labrador sea and the davis strait: early extension (220–150 ma, late triassic to late jurassic) increased extension (around 150 ma, kimmerid gian) regional rifting and dyke intrusion (140–133 ma, berriasian to valanginian/hauterivian) subsidence and sedimentation (130–120 ma, barre mian and aptian) faulting, sedimentation and magmatism (around 120– 100 ma, aptian–albian) late cretaceous subsidence, sedimentation and faulting (100–65 ma) palaeogene rifting and magmatism (about 62 ma). the last phase gave rise to the west greenland palaeogene volcanic province (fig. 2), with a volcanic • • • • • • • bulletin37.qxp_bulletin 37 26/01/17 13.59 side 16 17 succession up to 3 km thick (storey et al. 1998; larsen et al. 2016), occurring mainly onshore but extending offshore. onshore, volcanism associated with this pro vince commenced around 61–62 ma (chron 27n; latest danian to early selandian; nøhr-hansen et al. 2002), with most of the basalts being extruded in less than a million years (storey et al. 1998; larsen et al. 2009, 2016), followed by younger paleo cene volcanism in the nuussuaq basin. other volcanic basalts are con sidered to be of ypresian, late middle eocene and middle oli gocene age. according to larsen et al. (2009), the youngest volcanic rocks in west greenland are 28 ma (around the rupelian–chattian boundary; fig. 3). rol le (1985) noted that the jurassic and lower cretaceous alkaline intrusions in southern greenland are similar pe trographically to the intrusives of north-western new foundland, which he believed were associated with the formation of the labrador margin basins. north-eastern canada the oldest onshore cretaceous rocks on the canadian margin adjacent to the labrador–baffin seaway are the ford bight volcanics, which crop out along the la b rador coast. these rocks were described by king & mcmillan (1975) and dated by balkwill et al. (1990) as 129 ± 6 ma (valanginian – earliest aptian on the gradstein et al. 2012 timescale) and 145 ± 6 ma (ti thonian – earliest valanginian). these ages are broadly compatible with those given for the alexis formation (122 and 139 ma). there are no known rocks onshore labrador that equate with the bjarni formation, but such rocks do occur in eastern and north-eastern baffin island and on bylot island. from the cape dyer area of baffin island, burden & langille (1990) described sandstones, siltstones and coal beds that contain palynomorphs indicative of an aptian to early cenomanian age. outcrops of cretaceous and lower cenozoic rocks occur in the eclipse trough on bylot island and northeastern baffin island. these deposits, which show similarities with rocks on the labrador margin (figs 2, 3), were described by miall et al. (1980), mcwhae (1981) and miall (1986), who used lithostratigraphic unit names with type sections in the arctic islands to the north-west. the units are, from oldest to youngest, the hassel for mation (equivalent to the bjarni formation), the kanguk formation (equivalent in part to the markland formation) and the eureka sound group (fig. 3). the eureka sound group was named by troelsen (1950), reduced to formation status by tozer (1963), and raised back up to group status independently by miall (1986) and ricketts (1986). rather than using the eureka sound formation/group, sparkes (1989) and water field (1989) proposed an informal lithostratigraphic breakdown for equivalent rocks on bylot island. accor ding to these authors, the kanguk formation is unconformably overlain by the informal sermilik and bylot island formations, which are in part coeval. the top of the bylot island formation is unconformably overlain by the pond inlet formation, above which are the informal navy board and aktineq formations (fig. 3). this scheme was followed by harrison et al. (1999) and har rison et al. (2011). mcwhae (1981) considered the unconformity separating the hassel formation from the overlying kanguk formation to be the avalon unconformity and he named the hiatus between the kanguk and eureka sound group (as eureka sound formation) as the bylot unconformity. although not using the name bylot unconformity, miall (1986) also recognised an unnamed unconformity between the kanguk for mation and the eureka sound group. both of these authors indicated an early paleocene age for this unconformity. however, sparkes (1989) and waterfield (1989) dated the bylot unconformity as maastrichtian. similarly, harrison et al. (2011) gave a maastrichtian to danian age to the section extending from the sermilliq formation to the aqtineq formation (i.e. the former eureka sound formation/group), with an unconformity beneath these. as discussed further below, the number of unconformities and/or hiatuses is probably greater than the five specified by mcwhae et al. (1980). for example, our data seem to indicate two hiatuses associated with the maastrichtian–paleocene interval, one at the base of the cenozoic and one in the selandian. on baffin island, andrews et al. (1972) recorded marginal marine to lacustrine palaeogene sediments from the north-central region, and clarke & upton (1971) found terrestrial sediments north of cape dyer. burden & langille (1990) described fluvial and debrisflow deposits of danian age on the quqaluit and pad loping islands (figs 1, 2). presumably these rocks and the overlying cape dyer basalt represent the late rifting phase immediately before the opening of baffin bay. pedersen et al. (2002) considered the cape dyer basalt most likely to be time equivivalent to the lower part of the vaigat formation in the nuussuaq basin (fig. 3). bulletin37.qxp_bulletin 37 26/01/17 13.59 side 17 18 the results presented in this bulletin reflect data from analyses of offshore wells from both margins of the la b rador–baffin seaway; detailed taxonomic data are pre sented in the companion bulletin (fensome et al. 2016). the labrador margin wells are from both the hopedale basin (bjarni o-82, north leif i-05, ogmund e-72, roberval k-92, snorri j-90 and south labrador n-79) and the saglek basin (gilbert f-53, gjoa g-37, hekja o-71, karlsefni a-13, ralegh n-18, rut h-11 and skolp e-07). the wells on the west greenland margin are: hellefisk-1, ikermiut-1, kangâmiut-1, nukik-1, nu kik-2 and qulleq-1. also incorporated are some results from onshore west greenland cored boreholes and sur face sections, and from some shallow cored boreholes from western baffin bay. an event-stratigraphic approach is adopted as we consider this to provide greater detail and precision than an approach based on zonations. an event-stratigraphic approach also better facilitates incorporation of events from multiple disciplines; the events used here are exclusively derived from palynology. events can be originations of species, called first occurrences (fos), extinctions, called last occurrences (los), or occasionally peak occurrences, or acmes, of species. although some stratigraphic control was obtained from analysis of conventional cores from several la b rador margin wells and from some side-wall cores, the availability of and recovery from such samples were limited. it has been necessary therefore to rely mostly on ditch-cuttings samples, and hence on los for most of the events, as such samples usually contain downsection contamination (cavings). reworked specimens also occur in many samples, being most frequent in what we consider to be pliocene–pleistocene sections. piasecki (2003) noted that reworked species outnumbered in-situ species in his late pliocene interval in the qulleq-1 well. contamination from drilling mud was also a minor concern; for example, the presence of the dinocyst ovoidinium verrucosum in most of the ditchcuttings samples in hellefisk-1 and nukik-2 is interpreted to have resulted from contamination from cen o manian bentonite added during drilling. samples processed at gsca (bjarni o-82, gilbert f-53, karlsefni a-13, roberval k-92, rut h-11, snorri j-90, south labrador n-79) followed the procedure outlined in barss & williams (1973) and carried out from 1971 to 2007 (fig. 4). the first step after crushing the sample, if necessary, was the addition of 10% hydrocholoric acid to remove carbonates. after decanting the acid, the residue was washed and placed in hydrofluoric acid to remove the silicates. when the sample appeared to be broken down, the hydrofluoric acid was decanted and the sample washed several times. to remove any remaining fluorides, the sample was placed in concentrated hydrochloric acid, followed by washing. the fraction remaining was further concentrated using a heavy liquid such as zinc bromide. this was followed by oxidation in concentrated nitric acid or schulze solution, followed by careful washing to re move the oxidising agents, and treatment with ammonium hydroxide. the final phase before making the slides was sieving, a technique that concentrates the dinocysts and larger miospores. all the slides were mounted in elvacite. most residues were stained with safranin red or bismarck brown. although oxidation, if done at all, was always carefully controlled, most samples were processed at a time when the detrimental methodology mudstone, sandstone, limestone pre-acid preparation 10% hcl decant wash 3 times concentrated hf decant wash 3 times concentrated hcl decant wash 3 times centrifuge wash 3 times hn4oh wash until clear remove remaining minerals stain screen 20, 30, 38, 180 µm stain/make slides coal, peat znbr2 1.4 specific gravity differential centrifuge/ screen store cut for carbonisation studies store residue oxidise (hno3 or schulze solution) znbr2 2.0 specific gravity remove coarse material fig. 4. flow chart showing the palynological processing method used at the geological survey of canada (atlantic). bulletin37.qxp_bulletin 37 26/01/17 13.59 side 18 19 selective effect of oxidation procedures on peridinioids, and especially protoperidinioids, was not fully appreciated. hence the general scarcity of protoperidiniod cysts in our assemblages may be due to over-oxidation, although most samples stratigraphically predate times when protoperidinioids became common. samples from hellefisk-1, ikermiut-1, kangâmiut1, nukik-1, nukik-2, qulleq-1, north leif i-05, og mund e-72, skolp e-07, hekja o-71, ralegh n-18 and gjoa g-37 were processed at the geological survey of greenland (ggu, now geus). palynomorphs were extracted from 10–20 g of sediment from each sample by modified standard preparation techniques that in cluded treatment with hydrochloric (hcl) and hydrofluoric (hf) acids, sieving using a 20 μm nylon mesh and oxidation (3–10 minutes) with concentrated nitric acid (hno3), often followed by washing with a weak potassium hydroxide solution (koh). finally, palynomorphs were separated from coal particles and woody material in most samples using the separation method described by hansen & gudmundsson (1978) or by swirling. after each of the steps mentioned above, the organic residues were mounted in a solid medium (eukitt®) or in glycerine gel. the palynological slides were studied with transmitted light using a leitz dialux 22 microscope (no. 512 742/057691). dinocysts, acri tarchs and selected stratigraphically important spore and pollen (miospores) species were recorded from the sieved, oxidised or gravity-separated slides. approxi mately 100 specimens were counted whenever possible. using transmitted light microscopy, qualitative and quantitative analyses were undertaken of the palyno morphs present in each well (see appendices 2, 3); the stratigraphically important events are plotted on sum mary charts (figs 5, 6, in pocket). the los of species plotted in figs 5 and 6 were generally based on occurrences in several wells, thus allowing correlation be tween well sections. however, few of the taxa recorded on the events plot are present in all the wells, due in part to the variable duration of the hiatuses in the individual wells. a detailed chart has been generated for each well showing los and other key events, such as peak occurrencies and acmes (appendix 3). although broad con sistency has been sought, the information included in these charts is variable to some degree, dependent on the source of the data. those done by hn-h (helle fisk-1, ikermiut-1, kangâmiut-1, nukik-1, nukik-2, qulleq-1, north leif i-05, ogmund e-72, skolp e-07, hekja o-71, ralegh n-18 and gjoa g-37) present a zonation and events (predominantly los) that were published in sønderholm et al. (2003b), nøhr-hansen et al. (2000) and nøhr-hansen (2003, 2004a, b) for the aptian‒albian to priabonian of the west greenland and labrador margins of the labrador sea. some of hn-h’s plots in the present paper also show dinocyst species richness and azolla abundances. plots generated by glw and raf (bjarni o-82, gilbert f-53, karl sefni a-13, roberval k-92, rut h-11, snorri j-90 and south labrador n-79) show events and ages only; a zonation was not developed. in total, 187 bioevents are identified for the labra dor–baffin seaway; 169 of these represent the youngest or last occurrence (lo) of a taxon, the others represent peak occurrences or abundances. some events are coe val, so the 187 bioevents define 106 bioevent horizons. collectively, the bioevents are based on 177 taxa. also incorporated are 22 events (13 event horizons) established by piasecki (2003) for the neogene of off shore west greenland and 50 events (44 event horizons) established by pedersen & nøhr-hansen (2014) for albian to paleocene strata of the nuussuaq basin, west greenland (figs 5, 6 in pocket). most of the species are dinocysts but several are pollen and spores, and we also include a fungal spore peak and two azolla peaks. to promote understanding, communication and future consistent recognition of taxa encountered in this study, a companion bulletin presents the systematic taxonomy (fensome et al. 2016), describing and illustrating most of the species referenced in figs 5 and 6 and specifying location and repository information for the material used in this study. the timescale of gradstein et al. (2012) is adopted. detailed analyses and palaeoenvironmental curves for wells analysed in this study are presented in appendices 2 and 3; those dinocyst taxa whose occurrences are considered to have significance as palaeoenvironmental indicators are listed under the appropriate palaeoenvironment in table 1 (see p. 37). bulletin37.qxp_bulletin 37 26/01/17 13.59 side 19 20 palaeontological evidence used to interpret the evolution of the labrador–baffin seaway region has been derived primarily from the study of cretaceous–neo gene microfossil (including palynological) assemblages recovered from offshore wells on the labrador margin and on the west greenland margin. papers, reports and two theses on the palynomorphs from labrador margin and davis strait wells, all drilled in the 1970s and 1980s, have been produced by williams (1975), gradstein & williams (1976), barss et al. (1979), ioakim (1979), gradstein & williams (1981), wil liams (1986), bujak davies group (1987), bell (1989), wil liams et al. (1990), nøhr-hansen (2004a, b), nøhr-han sen (in sønderholm et al. 2003b), williams (2007a–f), ainsworth et al. (2014) and fensome (2015). fora miniferal studies were undertaken by grad stein & wil liams (1976), gradstein & srivastava (1980), grad stein & agterberg (1982), gradstein et al. (1994) and ains worth et al. (2014). williams (1986) is the only author to have presented a zonation for labrador–baffin seaway wells based on cenozoic spores and pollen. she analysed six wells: kangâmiut-1, hekja o-71, karlsefni a-13, herjolf m92, roberval k-92 and cartier d-70. based on these analyses, she recognised eight provisional spore and pollen zones of the following ages: early to middle paleo cene, middle to late paleocene, early to middle eocene, middle to late eocene, latest eocene/earliest oligocene, early oligocene, early to middle miocene, and middle to late miocene. the middle and upper oligocene was inferred to be absent in the wells. ioakim (1979) analysed the dinocysts from two hopedale basin wells, bjarni h-81 and freydis b-87. she erected 11 zones and three subzones, which collectively spanned the maastrichtian–priabonian. most of the zones were named after peridiniaceans, especially wetzelielloideans. ainsworth et al. (2014) undertook a detailed litho stratigraphic and biostratigraphic study of six labrador margin wells: herjolf m-92, north bjarni f-06, og mund e-72, pothurst p-19, roberval k-92 and snorri j-90, plus hare bay e-21 from offshore north-eastern newfoundland. they based their biostratigraphy on fora minifera and palynomorphs, with the latter being primarily dinocysts, and their palaeoenvironmental in terpretations exclusively on fluctuations in the for aminiferal assemblages. among their suggestions was a division of the markland formation into a lower mem ber of coniacian–maastrichtian age and an upper mem ber of selandian age, with the boundary between the two units equivalent to a base-cenozoic unconformity. according to ainsworth et al. (2014), the upper part of the markland formation was deposited in a predominately deep-water setting – a very different interpretation from that of balkwill & mcmillan (1990), who postulated a marginal marine to middle shelf environment. the re-interpretation by ainsworth et al. (2014) was based on present-day knowledge of deep-water agglutinated foraminifera. ainsworth et al. (2014) in cluded the gudrid member as part of the cartwright formation (instead of being parts of the markland, cartwright and kenamu for mations as used here), and considered it to be thane tian in age and deposited in a bathyal environment. based on their interpretation of the saglek basin’s pothurst p-19 well, ainsworth et al. (2014) regarded much of the ‘lower’ mokami for ma tion (chattian to miocene) and all of the ‘upper’ mio cene to be absent. this is at variance with other studies, which indicate that miocene sediments occur in this well. the interpretation of ainsworth et al. (2014) sup ports the observations of knutsen et al. (2012), how ever, who noted that off west greenland during the miocene there was uplift and erosion, which they related to hot-spot migration. a recent study of pothurst p19 (g.l. wil liams, unpublished data) demonstrated that most of the lower and all of the middle miocene are missing, but some upper miocene is present. palaeogene nannofossil biostratigraphy was de scribed for the gilbert f-53 and skolp e-07 wells by crux & gard (2004). in gilbert f-53, they recorded a thick palaeogene sequence succeeded by an indeterminate cenozoic succession. skolp e-07 contained a sec tion with mixed nannofossils of cretaceous and ceno zoic age overlain by cenozoic strata. crux & gard (2004) interpreted the lower section to be probably cenozoic with reworked cretaceous nannofossils. nan no fossil re covery was generally disappointing. on the west greenland margin, rolle (1985) published the stratigraphy of the five wells drilled in 1976 and 1977 (hellefisk-1, ikermiut-1, kangâmiut-1, nukik-1, nukik-2; figs 1, 2), with age control based on palynology. renewed interest in exploration of the west greenland margin in recent years has motivated new and detailed studies of the cretaceous to neogene previous palaeontological studies bulletin37.qxp_bulletin 37 26/01/17 13.59 side 20 21 palynology (e.g. nøhr-hansen 1996, 1997a–c, 2003; nøhr-hansen & dam 1997; dam et al. 1998b, 2009; nøhr-hansen et al. 2000, 2002; sønderholm et al. 2003b; piasecki 2003; pedersen & nøhr-hansen 2014). other microfossil groups utilised include nannofossils (nøhr-hansen et al. 2000; sheldon 2003) in the palaeo gene–neogene and foraminifera (nøhr-hansen et al. 2000; rasmussen et al. 2003; rasmussen & sheldon 2003), plus ostracods, radiolarians and diatoms (ras mussen et al. 2003). further information comes from numerous creta ceous–cenozoic outcrop sections from the nuussuaq basin in west greenland, a few onshore sections in labrador and on baffin and bylot islands, and from odp leg 105, hole 645 from offshore baffin island (fig. 2; kaminski et al. 1989a, b). invariably, the bio stratigraphy is based on microfossils including palynomorphs. results from analyses of the cuttings provided the basis for the plots of los of taxa (figs 5, 6, in pocket). peak occurrences or abundances are also included in the palaeogene. as noted previously, results are based on three areas: the hopedale basin on the labrador mar gin (bjarni o-82, north leif i-05, ogmund e-72, ro berval k-92, snorri j-90 and south labrador n-79), and the saglek basin, partly on the labrador margin (gilbert f-53, karlsefni a-13, rut h-11 and skolp e07) and partly off south-eastern baffin island (gjoa g37, hekja o-71, ralegh n-18). on the west green land margin, the wells are: hellefisk-1, ikermiut1, kangâ miut-1, nukik-1, nukik-2, and qulleq-1. figures 5 and 6 represent composite compilations that involve the merging of data from both the canadian and green land margins. one unexpected difficulty has been the identification of taxa common to both mar gins of the seaway. this may be because some taxa are endemic to one margin or even one area, and perhaps also because of the different latitudes of the overall location of wells on the two sides, the greenland wells being farther north. however, there was enough overlap to provide a detailed event-stratigraphic framework (figs 5, 6, in pocket), especially with the incorporation of information on known stratigraphic ranges of taxa in european sections (powell 1992; bujak 1994; williams et al. 1999, 2004) and from other wells and cored boreholes from offshore eastern canada and west green land (williams 1975; williams & brideaux 1975; williams & bujak 1977; barss et al. 1979; williams et al. 1990; sønderholm et al. 2003b; nøhr-hansen 2004a, b; fen some & williams 2005; fensome et al. 2008, 2009). lower cretaceous the oldest mesozoic rocks identified here in the la brador–baffin seaway are aptian (see below), and these are restricted to the labrador margin. however, h. nøhrhansen has also observed reworked late jurassic – early cretaceous taxa such as lithodinia sp., nelchinopsis kostro miensis, sirmiodinium grossii, and forms similar to gony aulacysta pectinigera/fastigiata from seabed samples col lected by geus (dalhoff et al. 2006). piasecki (2003) also reported late jurassic dinocysts in dredge samples from the west greenland margin. aptian gradstein & williams (1976), in their biostratigraphic study of labrador margin wells, defined a cerebro polle n ites mesozoicus assemblage of barremian–aptian age. in this study, it has not been possible to confirm the presence of barremian strata, but an aptian age is demonstrated on the labrador margin by the presence of the dinocyst tenua hystrix, which is the taxonomic senior synonym of cer bia tabulata (see fensome et al. 2016). tenua hystrix, misidentified as cyclonepheliium atta da li cum, was the index species for an early aptian zone in the scotian basin (williams 1975). duxbury (2001) placed the lo of cerbia tabulata at the top of the biostratigraphic results bulletin37.qxp_bulletin 37 26/01/17 13.59 side 21 22 aptian, which is at 113 ma on the gradstein et al. (2012) timescale. other studies (e.g. williams 2003b) have confirmed the lo of cerbia tabulata to be a consistent marker for the top of the aptian. also in the aptian of the roberval k-92 and south labrador n-79 wells is a form that we identify as pseu doceratium sp. wells with aptian rocks include bjarni o-82 (appen dix 3.3), roberval k-92 (appendix 3.2) and south labrador n-79 (appendix 3.4; figs 1, 7). palyno morphs found in the aptian rocks in our study include the miospores callialasporites dampieri, callialasporites obrutus, cerebropollenites mesozoicus, par vi saccites amplus, appendicisporites jansonii, cicatrico sis porites australiensis and klukisporites areolatus. ikermiut-1 qulleq-1 offshore west greeenland onshore west greeenland offshore eastern canada labrador margin, saglek basin gilbert f-53 gro#3 well nuussuaq umiivik-1, svartenhuk halvø & north nuussuaq, surface section skolp e-07bjarni o-82 south labrador n-79 snorri j-90 ogmund e-72 offshore eastern canada, labrador margin, hopedale basin 110 115 120 125 105 70 75 80 85 90 95 100 ma north leif i-05 roberval k-92 110 115 120 125 105 70 75 80 85 90 95 100 l at e c re ta ce ou s ea rl y c re ta ce ou s epochma maastrichtian campanian santonian coniacian turonian cenomanian albian aptian age fig. 7. cretaceous ages (highlighted in green) identified from palynomorph assemblages in wells from the hopedale basin of the labrador margin (north leif i-05, roberval k-92, bjarni o-82, south labrador n-79, snorri j-90 and ogmund e-72), the saglek basin of the labrador margin and davis strait (skolp e-07 and gilbert t-53), offshore west greenland (qulleq-1 and ikermiut-1), and a well (gro #3, nuussuaq), a borehole (umiivik-1) and two surface sections (svartenhuk halvø and north nuussuaq) from onshore west greenland. ages not coloured have not been identified. the timescale (ma) is from gradstein et al. (2012). bulletin37.qxp_bulletin 37 26/01/17 13.59 side 22 23 albian albian assemblages in labrador margin wells strongly reflect marginal marine to innermost neritic palaeoenvironments. dinocyst taxa with albian los include nyktericysta davisii, nyktericysta dictyphora, nyktericysta tripenta, oligosphaeridium albertense, vesperopsis longicornis and subtilisphaera perlucida. nyktericysta davisii was described from middle to upper albian brackishwater deposits in the western interior seaway of the usa by bint (1986). in offshore eastern canada, it has a consistent lo in the albian and other records of its occurrence in the late albian are from onshore east and west greenland and arctic canada (nøhr-hansen 1992, 1993a, 2008; macrae 1996; nøhr-hansen & mcintyre 1998) and the scotian mar gin (fensome et al. 2008). however, a recent study by pedersen & nøhr-hansen (2014) indicated that large forms of nyktericysta davisii and nyktericysta arachnion occur in the early cenomanian of onshore nuussuaq basin based on co-occurrence with the pollen species rugubi ve siculites multisaccus, which is indicative of an early cenomanian age according to singh (1983). nykteri cysta dictyphora is also characteristic of brackish-water deposits in the aptian–albian of china (mao shaozhi et al. 1999). senoniasphaera microreticulata oc curs in the albian of north leif i-05 (nøhr-hansen 2004b). miospores dominate most albian assemblages on the labrador margin. bujak davies group (1987) defined a parvisaccites amplus zone which they considered early albian. conventional cores from north leif i-05 con tain parvisaccites radiatus with rugubivesiculites rugosus, rugubivesiculites multiplex, rugubivesiculites reductus and rugubivesiculites convolutus, and conventional cores from bjarni o-82 contain rugubivesiculites rugosus. wil liams (1975) erected a rugubivesiculites rugosus as semblage subzone of late albian age for the scotian shelf and grand banks. singh (1983) placed the lo of rugubivesiculites rugosus in the cenomanian and nøhrhansen (in sønderholm et al. 2003b) recognised a ?late albian to ?cenomanian rugubivesiculites spp. inter val, based on a study of skolp e-07 and ogmund e-72. the cored interval in north leif i-05 and bjarni o-82 is thus no older than late albian. we have identified albian sediments in bjarni o-82 (appendix 3.3), north leif i-05 (appendix 3.1), ogmund e-72 (appendix 3.6), and skolp e-02 (ap pendix 3.7; figs 1, 7), and onshore in the nuussuaq basin. several taxa other than those plotted on the events chart (fig. 5, in pocket) have provided supporting data for age determinations in the aptian–albian, including pilosisporites verus, cicatricosisporites reticicatricosus, contignisporites glebulentus and alisporites grandis, all with los in the albian (williams 2007a). upper cretaceous cenomanian this study indicates that the cenomanian is present in only four wells on the canadian margin. however, there are some characteristic markers for this stage, including the dinocyst species kiokansium williamsii described by singh (1983) from the albian and ceno manian of alberta: in this study, it was found in north leif i-05 and bjarni o-82. the species was first recorded as cleistosphaeridium polypes subsp. a by williams (1975), who noted that its lo was in the ceno manian. other characteristic dinocyst taxa of the cenomanian are oligosphaeridium totum and odonto chitina ancala. williams et al. (1999) placed the lo of oligo sphae r idium totum at 96.2 ma, in the middle ce nomanian whereas fensome et al. (2008) indicated its lo to be earliest cenomanian. an expecially distinctive morphotype with an lo in the cenomanian is the pol len afropollis (fensome et al. 2009); cicatri cos i sporites mi nu ta estriatus has a cenomanian lo in og mund e-72. all the wells with cenomanian strata – north leif i-05 (appendix 3.1), ogmund e-72 (ap pendix 3.6) and skolp e-07 (appendix 3.7) – are on the labrador mar gin (figs 1, 7). the upper cenomanian in the nuus suaq basin is characterised by the fos of tri thy ro d i nium suspectum, isabelidinium magnum and cauve rid i nium membraniphorum (pedersen & nøhr-hansen 2014; fig. 5, in pocket). turonian–santonian the turonian appears to be absent from all offshore wells, except possibly lower turonian in north leif i05. because of uncertainty we have tended to define a turonian–coniacian, rather than a turonian, interval for the labrador wells. species of rugubivesi culites appear to have their los in the turonian, but may range higher. the lower turonian in the nuus suaq basin is characterised by the fo of hetero sphaeri dium difficile (nøhr-hansen 1997a; pedersen & nøhrhansen 2014). in the umiivik-1 borehole on svarten huk halvø (fig. 2), dam et al. (1998b) recognised an upper turonian interval iv and an uppermost turo bulletin37.qxp_bulletin 37 26/01/17 13.59 side 23 24 nian to lower coniacian interval iii; interval iv is charac terised by a heterosphaeridium difficile acme; he tero sphaeridium difficile and circulodinium distinctum are common in the upper part of interval iii and raphi dodinium fucatum has its acme in the upper part of interval iii. interval iii correlates in part with the early coniacian spinidinium echinoideum interval described by nøhr-hansen (1996). bujak davies group (1987) recognised a heterosphaeridium difficile zone, which they considered to be early campanian. the top of nøhr-hansen’s (1996) coniacian – early santonian heterosphaeridium difficile interval was defined by the lo of heterosphaeridium difficile at the top of the early santonian, citing as justification costa & davey (1992). williams et al. (2004) plotted the fo of hetero sphaeri dium difficile at about the early–middle turonian boun dary and its lo in the late coniacian in northern midlatitudes. in a study of some shallow cored boreholes from the baffin margin, maclean et al. (2014) considered heterosphaeridium difficile to have a range of late coniacian to early santonian, but noted that it was oc casionally found as high as the early campanian. nøhrhansen (2012) and radmacher et al. (2014) noted that the species occurred in turonian–coniacian sediments in the kangerlussuaq basin, southern east greenland and in the south-western barents sea. in this study, we have placed the lo of heterosphaeridium difficile at the top of the early santonian. support for this is provided by the lo of odontochitina porifera, which stover et al. (1996) placed in the early san tonian, although other authors (costa & davey 1992; williams et al. 2004) have extended the range of the species into the cam panian. neither heterosphaeridium difficile nor odon to chi tina porifera are found in the only two west green land wells with cre taceous rocks, qulleq-1 (ap pendix 3.14), and iker miut-1 (appendix 3.18), indicating that the oldest sediments in these wells are younger than early santonian. on the labrador margin, bjarni o-82 (ap pendix 3.3) and south labrador n-79 (appendix 3.4) appear to have coniacian sediments (figs 1, 7). campanian the lo of spongodinium (originally scriniodinium) obscurum is taken here to mark the santonian–cam panian boundary, a somewhat arbitrary decision. spon go dinium obscurum was erected by manum & cookson (1964), who considered the sample from graham is land, arctic canada, from which the species was first re covered, to be early late cretaceous in age. this was based on the occurrence of the species in what these authors considered to be the hassel formation, of late albian to early cenomanian age. however, felix & burbridge (1976) demonstrated that the sample was from the upper cenomanian – lower campanian kan guk formation, and specifically from the upper kan guk formation. this reassignment would indicate that the age of the graham island sample is santonian to early campanian, and that the lo of spongodinium obscurum is within this time interval. the base of the lower to middle campanian aquilla pollenites interval de cribed by nøhr-hansen (1996) from the nuussuaq basin is defined by the fo of aquillapollenites species. in the labrador margin wells, the following dinocyst species have their los at the campanian–maastrichtian boundary: odontochitina co stata, trichodinium castanea and xenascus ceratioides. stover et al. (1996) considered the lo of odontochitina costata to equate with the campanian–maastrichtian boundary but extended the lo of odontochitina operculata into the earliest maa strichtian. these authors also extended the lo of xe nas cus ceratioides to the top of the early maastrichtian, but placed the lo of tri chodinium castanea just above the base of the late campanian. williams et al. (2004) gave the lo of xe nascus ceratioides as just above the campanian–maa strichtian boundary in northern midlatitudes. trithyrodinium suspectum and raphidodinium fucatum have los within the later campanian. this range conforms with that in williams et al. (2004), who placed the los of both species in the late campanian, but considered raphidodinium fucatum to be the young er. nøhr-hansen (in sønderholm et al. 2003b) recognised an odontochitina operculata interval, which was regarded as being of late campanian age. one of the characteristic species was heterosphaeridium bellii (as heterosphaeridium heteracanthum), which has its lo in the latest campanian in the labrador–baffin sea way. heterosphaeridium bellii was described by rad macher et al. (2014), who stated that it had a common occurrence in the late campanian to early maa strichtian. rad macher et al. (2014) defined the top of the early cam panian on the lo of callaiosphaeridium asymmetricum, which accords well with our findings. wells with cam panian sections are roberval k-92 (ap pendix 3.2), bjarni o-82 (appendix 3.3), ogmund e-72 (appendix 3.6), skolp e-07 (appendix 3.7) and gilbert f-53 (appendix 3.9) on the labrador margin and qulleq-1 (appendix 3.14) and ikermiut-1 (ap pendix 3.18) on the west greenland margin (figs 1, 7). bulletin37.qxp_bulletin 37 26/01/17 13.59 side 24 25 maastrichtian palynomorph taxa with los in the maastrichtian include the dinocyst isabelidinium cretaceum. in a study of dinocysts from the campanian–paleocene of sey mour island and adjacent islands in antarctica, askin (1988) defined a zone 1 that was characterised by isabelidinium cretaceum. she considered the zone to be of late campanian age. bowman et al. (2012), based on updated data from seymour island, cited the age to be questionable late maastrichtian but recognised two younger maastrichtian zones. thus, the lo of isabeli dinium cretaceum can be considered as early late maa strichtian. the species is thus a key index species for determining whether part of the upper maa strichtian is missing in the labrador–baffin seaway. one stratigraphically significant pollen species is wodehousiea spinata, which occurs only in skolp e-07 on the labrador margin: its lo approximately equates with the maastrichtian–danian boundary. a dinocyst species with a similar lo is palynodinium grallator. nøhr-han sen (1996) described lower maastrichtian cerodi nium die belii and upper maastrichtian wode houseia spinata intervals from the nuussuaq basin on the basis of the fos of cerodinium diebelii and wode houseia spinata respectively, and nøhr-hansen (in sønderholm et al. 2003b) defined a palynodinium grallator interval for the uppermost maastrichtian. this was characterised by the los of palynodinium grallator and wodehouseia spinata. wil liams et al. (2004) placed the fo of palynodinium grallator in northern mid-latitudes within the late maa strichtian and its lo in the earliest danian. palynodinium grallator has been re corded only from onshore west greenland and from some labrador margin wells. disphaerogena carposphaeropsis, a dinocyst species that occurs in the latest maastrichtian of some labrador margin wells and in the nuussuaq basin (nøhrhansen & dam 1997; dam et al. 1998c) also appears to be a good index species. other such species in the maastrichtian include impagidinium victorianum, isa be li d i nium cretaceum, laciniadinium arcticum and spi ni ferites scabrosus. confirmation of the maastrichtian age for the lo of impagidinium victorianum comes from combined palynological and micropalaeontological analyses of south labrador n-79 by bujak davies group (1987). in an interval designated as maa strich tian based on foraminiferal data, they recognised an early maastrichtian subzone for a taxon they identified as impagidinium #ll. this taxon is probably conspecific with the species identified here as impagidinium victorianum and the species that nøhr-hansen (1996) identified from the upper maastrichtian in the nuus suaq basin as impagidinium sp. cf. i dispertitum. maa strich tian strata occur in several wells on the labrador margin (figs 1, 7), including north leif i-05 (ap pendix 3.1), roberval k-92 (appendix 3.2), bjarni o82 (appendix 3.3), south labrador n-79 (appendix 3.4), ogmund e-72 (appendix 3.6), skolp e-07 (ap pendix 3.7) and gilbert f-53 (appendix 3.9). several taxa other than those plotted on the events chart (fig. 5, in pocket) have provided supporting data for age determinations in the late cretaceous, including senoniasphaera protrusa (lo coniacian; williams 2007b), surculosphaeridium longifurcatum (lo conia cian), microdinium ornatum (lo coniacian), xenascus sar jean tii (lo coniacian), chatangiella ditissima (lo early campanian) and stiphrosphaeridium dictyophorum (lo maastrichtian). triprojectate pollen are also potentially useful (see braman 2013), but in the sec tions studied their occurrence was too sporadic and often clearly reworked. palaeogene danian–selandian some major hiatuses occur in the palaeogene sections on both margins of the labrador–baffin seaway. the oldest is in the danian, which seems to be missing or incomplete in all the wells, with the exception of south labrador n-79. the presence of this stage is indicated by the occurrences of the dinocyst species cerodinium diebelii, phelodinium kozlowskii, spongodinium delitiense, senoniasphaera inornata, tanyosphaeridium xan thi opy x ides and trithyrodinium evittii. according to wil liams et al. (2004), senoniasphaera inornata is restricted to the da n ian, with its fo near the base and its lo at about 63 ma in northern mid-latitudes. the same authors placed the lo of spongodinium delitiense at about 64 ma in north ern hemisphere mid-latitudes. the lower danian tri thy ro dinium evittii zone of nøhr-hansen et al. (2002) was recognised in north leif i-05 (nøhr-hansen 2004b) and in ogmund e-72 and skolp e-07 (nøhr-hansen in sønderholm et al. 2003b), indicating a late danian – thanetian hiatus in north leif i-05 and ogmund e-72 and a late danian – ypresian unconformity/hiatus in skolp e-07. williams (2007b) also concluded that in the south labrador n-79 well, the lo of trithyrodinium evittii is within the danian. we place the lo of cero dinium diebelii at the danian–selandian boundary, based on the records from bjarni o-82 (williams 2007a), bulletin37.qxp_bulletin 37 26/01/17 13.59 side 25 26 south labrador n-79 (wil liams 2007b) and snorri j-90 (williams 2007c; appendix 3.5). this is in slight conflict with williams et al. (2004), who placed its lo at about 60 ma in northern hemisphere mid-latitudes, within the early selandian. cerodinium diebelii has been recorded from the selandian in nukik-2 (nøhr-hansen 2003; ap pen dix 3.16), but this may represent reworking. with the exception of this probably reworked record in nukik-2, none of the index species above occur in the west greenland margin wells, indicating that da n ian strata are generally absent. however, the presence of cerodinium kangiliense and senegalinium iterlaaense in the lower part of nukik-2 may indicate a late danian or early selandian age (fig. 8) according to nøhr-hansen & heilmann-clausen (2000). in the labrador margin and davies strait wells, the danian is generally incomplete or missing, with the exception of south labrador n-79 (appendix 3.4; fig. 8). nøhr-hansen et al. (2002) erected five dinocyst zones, of which three are correlated to nannofossil zones, for the danian sedimen tary succession underlying the radiometrically dated se landian ba salts in the nuussuaq basin. dinocyst species with an lo within or at the end of the selandian include palaeoperidinium pyrophorum, spini dinium echinoideum and palaeocystodinium bulliforme. nøhr-hansen (2003) recognised a late danian and three late thanetian intervals on the west greenland margin. however, nøhr-hansen (2003) was following a twofold division of the paleocene in which selandian is represented by the lower part of thanetian (see powell 1992 and mudge & bujak 1996). the three thanetian intervals, from oldest to youngest are: the palaeoperidi nium pyrophorum interval (p4), correlating with latest se landian; the areoligera interval (p5), correlating with lower thanetian; and the apectodinium spp. acme interval (p6), correlating with upper thanetian (equivalences according to mudge & bujak 2001). although not using the three-fold division of the paleocene, nøhr-hansen (2003) recorded several of the species that appear to characterise a selandian age. the top of his palaeoperidinium pyrophorum (p4) interval is marked by the los of palaeo peridi nium pyrophorum and palaeocystodinium bulliforme. wil liams et al. (2004) placed the los of palaeo cystodinium bulliforme and palaeoperidinium pyrophorum close to the selandian–thanetian boundary. thus it seems reasonable to equate nøhr-hansen’s p4 interval with the selandian. in the present study, the threefold division is adopted (following gradstein et al. 2012) and the palaeoperid inium pyrophorum interval (p4) is referred to the late selandian for the west greenland wells (figs 1, 8), viz. nukik-2 (appendix 3.16), kangâmiut-1 (appendix 3.17), ikermiut-1 (appendix 3.18) and hellefisk-1 (ap pendix 3.19). nøhr-hansen et al. (2002) defined the base of their late danian palaeocystodinium bulliforme zone by the fo of palaeocystodinium bulliforme; however they did not record the lo of palaeocystodinium bulliforme. po w ell & brinkhuis (in gradstein et al. 2012) placed the lo of palaeocystodinium bulliforme at the se landian–thane tian boundary. based on the co-occurrent los of palaeo peridinium pyrophrum and palaeo cystodinium bulli forme, we place the latter lo at the selandian–thanetian boundary. thanetian the thanetian is characterised by maximum abundances of areoligera gippingensis and glaphyrocysta di varicata (referred to here as the areoligera gippingensis complex). the upper part of the areoligera gippingensis interval (p5) of nøhr-hansen (2003) must equate in part with the thanetian. high abundances of areoligera gippingensis complex cysts in a number of the wells – such as gilbert f-53, gjoa g-37, hekja o-71, helle fisk-1, ikermiut-1, kangâmiut-1, karlsefni a-13, nu kik-1, nukik-2 and qulleq-1 – allow for correlation with the type thanetian of southern england (powell et al. 1996) and with the skua e-41 (williams 2003a) and terra nova k-18 wells (williams 2003b) on the grand banks, offshore newfoundland. one species with an lo at the thanetian–ypresian boundary is cerodinium glabrum following immediately above the lo of cero dinium speciosum. gradstein & williams (1976) and bujak davies group (1987) both defined a cerodinium facing page: fig. 8. cenozoic ages (highlighted in brown for the palaeogene and yellow for the neogene) identified from palynomorph assemblages in wells from the hopedale basin of the labrador margin (north leif i-05, roberval k-92, bjarni o-82, south labrador n-79, snorri j-90 and ogmund e-72), the saglek basin of the labrador margin and davis strait (skolp e-07, gilbert f-53, rut h-11, hekja o-71, ralegh n-18 and gjoa f-37), offshore west greenland (qulleq-1, nukik-1, nukik-2, kangâmiut-1, ikermiut1 and hellefisk-1) and a well onshore west greenland (gro #3, nuussuaq) and north nuussuaq surface sections. ages not coloured have not been identified. the timescale (ma) is from gradstein et al. (2012). td: total depth. bulletin37.qxp_bulletin 37 26/01/17 13.59 side 26 27 o ff sh o re e as te rn c an ad a, l ab ra d o r m ar gi n , h o p ed al e b as in h el le fis k1 t d i n a lb ia n t d i n p re ca m b ri an b as em en t t d i n p re ca m b ri an b as em en t t d i n b as em en t t d i n b as al t t d i n b as al t t d i n b as al t t d i n b as al t t d i n d an ia n t d i n b as em en t t d i n b as em en t t d in u p p er sa n to n ia n t d i n a p ti an t d i n c am p an ia n sh al es t d i n l o w er c am p an ia n t d i n "d ia b as e in tr u si o n " ? ? ?? t d i n a p ti an t d i n p al ae o zo ic t d i n a p ti an t d (g r o # 3 ) in u p p er c re ta ce o u s o ff sh o re e as te rn c an ad a, l ab ra d o r m ar gi n , s ag le k b as in o ff sh o re a n d o n sh o re w es t g re en la n d n u ki k1 n u ki k2 k an gâ m iu t1 ik er m iu t1 g r o # 3 a n d n . n u u ss u aq o u tc ro p q u lle q -1 g jo a f3 7 r al eg h n -1 8 h ek ja o -7 1 k ar ls ef n i a -1 3 r u t h -1 1 sk o lp e -0 7 g ilb er t f5 3 o gm u n d e -7 2 sn o rr i j9 0 so u th l ab ra d o r n -7 9 b ja rn i o -8 2 n o rt h l ei f i0 5 r o b er va l k -9 2 1 0 1 5 2 0 2 5 3 0 3 5 4 0 4 5 5 0 5 5 6 0 6 55 p le is . plio. holoc. miocene oligocene eocene paleocene l e l m e l e l m e l e palaeogeneneogene epoch age period ma d an ia n se la n d ia n t h an et ia n y p re si an l u te ti an b ar to n ia n p ri ab o n ia n r u p el ia n c h at ti an a q u it an ia n b u rd ig al ia n l an gh ia n se rr av al lia n to rt o n ia n m es si n ia n z an cl ea n p ia ce n zi an g el as ia n bulletin37.qxp_bulletin 37 26/01/17 13.59 side 27 28 speciosum zone, which they considered to be late paleo cene. similarily, in his study of the north leif i-05 well, nøhr-hansen (2004b) considered the lo of cero dinium glabrum (as cerodinium speciosum subsp. gla brum) to be in the late thanetian. two index species that have their los near the base of the thanetian according to williams et al. (2004) are alisocysta cir cumtabulata and alisocysta margarita. these two species occur in several of the west greenland margin, davis strait and labrador margin wells. for example, aliso cysta margarita has been recorded from gilbert f-53, hekja o-71, gjoa g-37, north leif i-05, ogmund e72, ralegh n-18, snorri j-90 and south labrador n79 on the canadian margin and hellefisk-1, iker miut-1, kangâmiut-1, nujkik-1, nukik-2, and qulleq-2 on the west greenland margin. the uppermost thanetian is characterised by maximum abundances of apectodinium spp. and the range of axiodinium augustum (williams et al. 2015), and equates with the interval p6 of nøhr-hansen (2003). high abundances of apectodinium spp. and the pre sence of axiodinium augustum were recorded in several wells, e.g. north leif i-05, south labrador n-79, snorri j-90, ogmund e-72, rut h-11, gjoa g-37, kangâmiut-1 and ikermiut-1. ypresian in the labrador–baffin seaway, some of the richest dinocyst assemblages occur in ypresian sediments. the top of the stage is considered here to be marked by the frequent but never abundant occurrence of the freshwater fern azolla. brinkhuis et al. (2006) stated that azolla is abundant in basal middle eocene marine sediments of nordic seas. these authors placed the onset of this phase at c. 49 ma and the termination at c. 48.3 ma, probably based on the gradstein et al. (2004) timescale. barke et al. (2012) agreed with brinkhuis et al. (2006) in considering the azolla blooms to occur at 49 ma and also regarded the age as early middle eocene. brinkhuis et al. (2006) postulated that the high concentrations of azolla in an acex core taken on the lomonosov ridge reflected in-situ growth. brinkhuis et al.(2006) also listed wells from the labrador–baffin seaway (karlsefni a13 in the saglek basin and bjarni o-82, north leif i-05 and snorri j-90 in the hopedale basin) as having lowermost middle eocene abundances of azolla. in this study, these wells were not found to contain notable numbers of azolla. some other wells, all in the northern saglek basin, had greater numbers: hekja o-71, ralegh n-18 and gjoa g-37 (appendices 3.11–3.13). even in these wells, however, azolla was not recorded in high concentrations, more suggestive of transported material than in-situ growth (see further discussion, p. 42–43). immediately below the azolla occurrences are the los of the dinocyst species achilleodinium biformoides, diphyes brevispinum and piladinium columna. michoux (1988) described piladinium (as charlesdowniea) co lumna from upper ypresian rocks of south-western france, stating that nannofossil analyses indicated an np13 zone age. nøhr-hansen (2003) designated a piladinium (as charlesdowniea) columna interval of late ypresian age, whose top is marked by the lo of the zonal species. bujak (1994) plotted the lo of both pilodinium (as charlesdowniea) columna and diphyes brevispinum within the np13 zone. in the norwegian– greenland sea, eldrett et al. (2004) calibrated the lo of piladinium columna with chron 22n and within nannoplankton zone np14a, at the top of the ypresian; this is somewhat later than other records, but is in close accordance with this study. piladinium columna is wide spread in the labrador–baffin seaway, occurring in the following wells: gilbert f-53, gjoa g-37, north leif i-05, ogmund e-72, ralegh n-18, snorri j-90 and south labrador n-79 on the canadian margin, and kangâmiut-1, nukik-2 and qulleq-1 on the west greenland margin (fig. 1). species of apectodinium occur throughout the ypre s ian. they include apectodinium parvum, with an lo immediately above the thanetian–ypresian bound ary, and apectodinium homomorphum, which has a peak abun dance at about 54 ma. species of apectodinium are found in a number of wells, including gjoa g-37, gil bert f-53, hekja o-71, karlsefni a-13, north leif i-05, ogmund e-72, falegh n-18, rut h-11, skolp e-07 and south labrador n-79 on the canadian margin, and helle fisk-1, ikermiut-1 and kangâmiut-1 on the west greenland margin. the ypresian is informally subdivided here into early and late, with the boundary based on the los of petalo dinium condylos, evittosphaerula foraminosa and scale no dinium scalenum. supporting evidence comes from bujak (1994), who plotted the consistent occurrence of petalo dinium condylos within the nannofossil np12 zone. oc curring above the lo of petalodinium condylos is the lo of ginginodinium? flexidentatum. this species appears to equate with trinovantedinium #la of bujak davies group (1987). these authors defined a trino vante di nium #la zone, which they considered to be late ypresian based on foraminiferal data in south la b rador n-79. ginginodinium? flexidentatum is present in the ypres ian bulletin37.qxp_bulletin 37 26/01/17 13.59 side 28 29 in a number of wells, both on the eastern and western margins of the labrador–baffin sea way. these include gilbert f-53, hekja o-71, karl se fni a-13, north leif i05, ogmund e-72, rut h-11, snorri j-90 and south labrador n-79 on the canadian margin, and ikermiut-1 and nukik-2 on the west greenland margin (fig. 1). one of the characteristic features of the ypresian and succeeding lutetian is the occurrence of fungal spore peaks in some of the wells. selected fungal taxa are shown in fensome et al. (2016, plate 20, figs 12–20). lutetian although much or all of the lutetian seems to be absent in many of the wells, especially on the west greenland margin (fig. 8), its top seems to be clearly defined by the lo of diphyes colligerum. bujak (1994) placed the last consistent lo of this species at the lutetian–bar tonian boundary. eldrett et al. (2004) considered the lo of diphyes colligerum to approximate with the lutetian–bartonian boundary in the norwegian–green land sea. we follow eldrett et al. (2004) in placing the lo of diphyes colligerum close to the lutetian–bar tonian boundary; this is in contrast to nøhr-han sen (2004b), who recorded the lo of diphyes collige rum at the top of the ypresian in north leif i-05. another important index species having its lo at the top of the lutetian is alterbidinium? bicellulum, originally de scrib ed by islam (1983) from the earnley for mation of the middle eocene bracklesham group of southern eng land; the age of the bracklesham group is latest ypresian – lutetian. nøhr-hansen (2003) defined the top of his late lutetian interval on the lo of four taxa, one of which was alterbidinium? cf. bicellulum. recognition of the earliest lutetian is facilitated by the los of eatonicysta ursulae, cordosphaeridium gracile and hystrichosphaeridium tubiferum. bujak (1994) placed the lo of eatonicysta ursulae just above the base of the lutetian, in the nannofossil np14 zone. he signified the importance of this species by designating an eatonicysta ursulae zone. eldrett et al. (2004) developed a refined biostratigraphy of a dsdp site and two odp sites in the norwegian–greenland sea based on magnetostratigraphic calibration. these workers determined that the lo of eatonicysta ursulae was close to the top of the nannofossil np14a zone, broadly confirming the findings of bujak (1994). this species is found in five wells: kangâmiut-1 and nukik-2 on the west greenland margin and gjoa g-37, snorri j-90 and south labrador n-79 on the canadian margin. bujak (1994) also considered the lo of hystricho sphaeri dium tubiferum to be in the lutetian, but in the lower part of the succeeding nannofossil np15 zone. other los within the early lutetian include those of diphyes ficusoides, stichodinium lineidentatum and so p hismatia tenui vir gula. some of the events in the late lutetian are the los of dapsilidinium pseudoinsertum and glaphyrocysta vicina and the peak of trithyro di n ium? conservatum. bartonian the bartonian is absent or condensed in the greenland wells but more developed on the canadian margin (fig. 8). following bujak (1994), the top of this stage is de fined at the lo of the dinocyst rhombodinium porosum. dinocysts with their los in the bartonian include cerebrocysta bartonensis, chiropteridium gilbertii, and ho mo tryblium tenuispinosum. chiropteridium gilbertii occurs only in wells on the labrador margin (williams 2007e, under the informal name hystrichokolpoma ‘gil bertii’). pollen provide much of the stratigraphic control for recognition of bartonian strata in the labrador–baffin seaway. these include the los of ex tra tri poropollenites spp., corsinipollenites oculusnoctis, pi stilli pollenites macgregorii and cicatricososporites eocenicus. azolla spp. also occur for the last time. priabonian priabonian sediments have been identified in only one west greenland margin well (hellefisk-1) but are pre sent in most of the labrador margin and davis strait wells. the key species for defining the top of this stage are areosphaeridium diktyoplokum, lentinia serrata and phthanoperidinium multispinum. powell & brinkhuis (in gradstein et al. 2012) placed the lo of areo sphae r i dium diktyoplokum at the priabonian–rupelian bound ary. unfortunately, this species is found only in helle fisk-1 on the west greenland margin. the los of len tinia serrata and phthanoperidinium multispinum ap pear to define the top of the priabonian in some grand banks wells (e.g. williams 2003a, b). support for the lo of lentinia serrata is provided by williams et al. (2004), who placed it 0.2 million years above the pria bonian–rupelian boundary. lentinia serrata is found in gilbert f-53, north leif i-05, ralegh n-18 and south labrador n-79 on the canadian margin. other lo events in the priabonian include those of the dinocysts glaphyrocysta texta, schematophora speciosa, bulletin37.qxp_bulletin 37 26/01/17 13.59 side 29 30 cordosphaeridium funiculatum, phthanoperidinium stockmansii, rhombodinium draco and heteraulacacysta po rosa. williams et al. (2004) placed the los of sche ma tophora speciosa and cordosphaeridium funiculatum within the priabonian but, in contrast to the observations in this study, considered rhombodinium draco to range well into the rupelian. most of the above species occur only in one well, all on the canadian margin. another distinctive priabonian marker is the spore cicatricosisporites ornatus, which occurs in snorri j-90. rupelian strata of rupelian age were not encountered in any of the west greenland margin wells, but rocks of this age are widespread on the canadian margin, being present in bjarni o-81 (appendix 3.3), gilbert f-53 (appendix 3.9), hekja o-71 (appendix 3.11), karlsefni a-13 (appendix 3.8), north leif i-05 (appendix 3.1), ralegh n-18 (appendix 3.12) and south labrador n-79 (ap pendix 3.4; figs 1, 8). the top of the stage is placed at the lo of the two dinocyst species enneadocysta magna and licracysta semicirculata (fig. 6, in pocket). enneadocysta magna was described by fensome et al. (2007) from oli go cene sections in wells drilled on the grand banks. wil liams (2003a) used the then-informal name enneadocysta magna when he noted that it ap peared to be a consistent rupelian marker species in some grand banks wells. licracysta (as areoligera) semicirculata was described by morgenroth (1966) from the middle oligocene. wil liams et al. (2004) placed the lo of licracysta (as areo ligera) semicirculata just below the rupelian–chattian boundary. this seems to accord well with the findings of this study. other dinocyst species that help to delineate the rupelian are licracysta corymbus, phthanoperidinium coreoides and apteodinium au stra liense. miospores with los in the rupelian include zli visporis spp. and peri poropollenites sp. williams (2007b, e). chattian the chattian is difficult to delineate, primarily because it is partly or completely absent in all wells apart from hekja o-71 and ralegh n-18. we have been able to determine it based only on the los of two dinocyst species, chiropteridium galea and deflandrea phosphor itica. williams et al. (2004) considered that both these species extended into the early aquitanian in northern mid-latitudes, but for the purposes of this study the los of these species are considered to mark the chat tian–aquitanian boundary. neogene miocene dinocyst assemblages in the neogene of the labrador– baffin seaway are sparse and preservation is generally poor. both factors, which reflect the coarse clastics that predominate in this interval, may account for some gaps in our interpretations of ages, including the failure to define the aquitanian, the oldest stage of the mio cene. but there is another possibility – that sediments of aquitanian age are absent in a number of the wells (fig. 8). another difficulty in determining ages in the neo gene sections is the abundance of reworked palyno morphs, both cretaceous and palaeogene. how ever, there are some taxa that facilitate recognition of the early miocene. these are primarily the dinocyst cor dosphaer i dium cantharellus and the spore osmunda ci d ites wellmannii. williams et al. (2004) placed the lo of cor do sphaeridium cantharellus in northern hemi sphere mid-latitudes at about 19.5 ma, within the burdigalian. middle miocene strata are indicated from several wells, though it has not been possible to differentiate the langhian as all the index taxa appear to extend up into the serravallian. these taxa include the pollen cary apollenites spp. and tiliaepollenites crassipites, and the dinocysts cleistosphaeridium diversispinosm, apteo di nium spiridoides and cannosphaeropsis passio (fig. 6, in pocket). tiliaepollenites crassipites was previously de scribed as bombacacidites sp. a by williams (1975) and wil liams & bujak (1977). williams (1975) considered the lo of this species to be middle miocene in scotian margin wells. this age was slightly modified in wil liams & bu jak (1977), who recorded the species from the labrador margin operculodinium centrocarpum as semblage that they considered provisionally to be mid dle to late mio cene. williams (2007e) agreed with williams (1975) in assigning a middle miocene age. cannosphaeropsis pas sio, first recorded as nemato sphae ropsis sp. a by wil liams & brideaux (1975) from the upper miocene of the grand banks, was described by de verteuil & norris (1996) from the upper middle miocene of maryland. piasecki (2003) placed the lo of this species at the serravallian–tortonian boundary in qulleq-l, an age confirmed by williams et al. (2004), who likewise plotted its lo at the serravallian–tor tonian boundary in northern hemisphere mid-latitudes and equatorial bulletin37.qxp_bulletin 37 26/01/17 13.59 side 30 31 low-latitudes. powell & brinkhuis (in gradstein et al. 2012) considered the lo of apteo di nium spiridoides to be at the top of the early miocene, which conflicts with the range documented here. like wise, head et al. (1989a) recorded the species from the lower miocene of odp hole 645e, baffin bay (figs 1, 2), but did not extend its range into the middle mio cene. williams et al. (1999) plotted the lo of apteodi nium spiridoides within the serravallian, however, which corresponds more closely with our data. cleistosphae ri dium diversispinosum is generally an abundant species worldwide, but is not common in samples from the offshore wells in the labrador–baffin seaway. head et al. (1989a) noted, however, that the closely related clei stosphae r idium (as systematophora) ancyreum was com mon in samples from the early and middle miocene of odp hole 645e, baffin bay (figs 1, 2). another related species is cleistosphaeridium placacanthum, which schreck et al. (2012) considered to have an lo at 10.6 ma (early tortonian) in the dinocyst assemblages from an almost continuous middle miocene to pliocene section in odp site 907a in the iceland sea. williams et al. (1999) plotted the lo of cleistos phae ridium diversispinosum at the serravallian–tortonian boundary. upper miocene sediments are more prevalent in the west greenland wells than in those on the canadian side, with the qulleq-1 well appearing to have the most complete section (piasecki 2003; appendix 3.14). los occurring within the tortonian include mini sphae r idium latirictum, (formerly cordosphaer idium minimum), hy stric hokolpoma rigaudiae, palaeocystodinium gol zo wense, spiniferites pseudofurcatus, operculodinium janduchenei, operculodinium giganteum, operculo di nium piaseckii, ed wardsiella sexispinosa and labyrintho dinium truncatum. in this study, the boundary between the two late miocene stages, tortonian and messinian, is marked by the los of three dinocyst species tuber culodinium vancampoae (at the boundary), and palaeo cystodinium golzowense and operculodinium cen trocarpum in the late tortonian. tuberculodinium vancampoae, described from the pleistocene by rossignol (1962) is a cyst of the extant genus pyrophacus. head et al. (1989a) recorded tuber culodinium vancampoae from only one sample in the questionable early late miocene of hole 645e. this taxon is found in modern sediments, in warm lower latitudes. it is presumed that changing climatic conditions, as noted below, resulted in its migration from higher to lower latitudes toward the end of the miocene, although there were warming trends in the pliocene–pleistocene. according to powell & brinkhuis (in gradstein et al. 2012), the lo of palaeocystodinium golzowense is within the later tortonian. head et al. (1989a) found that pa laeo cystodinium golzowense was common in the early and middle miocene of odp hole 545e, but became rare up section, in what these authors considered to be early late miocene. two dinocyst species, spiniferites pseudofurcatus and dapsilidinium pastielsii have regional los within the tortonian. head et al. (1989a) recorded rare but consistent occurrences of the closely related dapsilidinium pseu do colligerum from the middle miocene in odp hole 645e, baffin bay, possibly extending into the early late miocene. although mertens et al. (2014) treated dapsili di nium pseudocolligerum as a taxonomic junior synonym of dapsilidinium pastielsii, we prefer to retain dapsili dinium pseudocolligerum, which appears to have a restricted stratigraphic range in the labrador–baffin seaway. head & westphal (1999) noted that the latitudinal occurrences of this species contracted in the late miocene and pliocene, due to the cooling of the north atlantic in the late miocene and the evolution of the cold labrador current. thus climatic changes and oceanic currents rather than extinction explain the disappearance of the genus dapsilidinium and the species tuberculodinium van campoae from the labrador–baffin seaway in the late miocene. within the messinian, species of the pollen quer coidites are common but do not appear to range above the miocene. spiniferites ovatus, described from the late mio cene of japan, also has its lo at the top of this stage. pliocene–pleistocene we have not been able to differentiate the pliocene and pleistocene in the wells analysed for this bulletin, but piasecki (2003) made some determinations based on analysis of the qulleq-1 well (appendix 3.14). he plotted the lo of habibacysta tectata immediately above the base of the zanclean. this may be too high, based on the findings of schreck et al. (2012), who placed its lo in the middle langhian in odp hole 907a in the iceland sea. a probable explanation for this discrepancy is that the specimens from qulleq-1 are reworked. other spe cies recorded by piasecki (2003) as having their los in the zanclean include barssidinium graminosum, inver to cysta lacrymosa, selenopemphix nephroides, seleno pemp hix bre vi spinosa, and reticulatosphaera actinocoronata. pia se c ki (2003) also distinguished the piacenzian–gela sian bound ary based on the lo of cymatiosphaera invaginata. in our analyses, we recognised the pliocene–plei sto cene based on the presence of the los of the pollen taxa, bulletin37.qxp_bulletin 37 26/01/17 13.59 side 31 32 zonalapollenites igniculus, graminidites sp. a of williams & brideaux (1975), and compositoipollenites sp. b of williams & brideaux (1975). williams (1975) defined an artemisia–taraxacum zone, which he con sidered to be pliocene–pleistocene. furthermore, wil liams & bu jak (1977), in a study of four labrador margin wells, defined a zonalapollenites (as tsugae polle nites) igniculus assemblage of pliocene–pleistocene age. discussion cretaceous according to umpleby (1979), the oldest mesozoic rocks beneath the labrador sea are of barremian age. it has not been possible to confirm a barremian age in this study, but the presence of aptian rocks has been determined in the south labrador n-79, roberval k-92, snorri j-90, and bjarni o-82 wells (figs 1, 2, 7). questionable up per aptian ‒ albian rocks occur in og mund e-72 and north leif i-05. all of these wells are in the hopedale basin. in the saglek basin, albian–cenomanian rocks are encountered in the lowermost 475 m of skolp e-07. question able cenomanian strata are present in bjarni o-82, north leif i-05 and og mund e-72. in the bjarni o-82 well, a major hiatus is identified, with much of the albian and cenomanian missing. this is equivalent to the avalon unconformity (fig. 3) identified by mcwhae et al. (1980) as a regional unconformity in the labrador–baffin seaway. we recognise another hiatus in several wells: in south labrador n-79, aptian strata are overlain by coniacian rocks; in ogmund e-72, the cenomanian is overlain by campanian; in north leif i-05, ceno manian–turonian is overlain by maastrichtian; and in skolp e-07, albian– cenomanian strata are overlain by campanian strata. thus, much of the upper cre taceous section appears to be missing in the hopedale and saglek basins. these observations accord with those of balkwill et al. (1990), who noted that the cen o manian–santonian succession is condensed, discontinuous and sometimes absent on the labrador margin. (fig. 7). in offshore west green land, the oldest drilled cretaceous rocks are in the qul leq1 well, where upper santonian rocks were en coun tered in the lowermost c. 300 m of the well. the continuation of the rift phase during the latest cretaceous is represented by thicker, more continuous sequences, although significant gaps in the succession still occur. in parts of the hopedale basin, such as at bjarni o-82 and ogmund e-72, the campanian is incomplete, with the lower part presumed missing. at north leif i-05, the campanian is absent. this con trasts with the situation in skolp e-07, where the campanian is about 1050 m thick. in offshore west greenland, lower campanian sediments occur only in the ikermiut-1 and qulleq-1 wells, with the upper cam panian and maastrichtian absent. reworked specimens of albian, turonian and late maastrichtian taxa are pre sent in the lowermost part of the kangâmiut-1 well, which is dated as paleocene. the upper cretaceous – palaeogene sequence on the labrador margin is variably developed. in the skolp e07 well, for example, the upper maastrichtian succes sion appears to be overlain by a thin succession of possibly danian age, which is in turn overlain by ypresian strata. other wells such as north leif i-05 and south labrador n-79 in the hopedale basin have more com plete sections, including maastrichtian rocks in the for mer and danian rocks in the latter (nøhr-hansen 2004b). onshore, in the nuussuaq basin, aptian? to al bian strata are the oldest mesozoic rocks overlying basement (dam et al. 2009; pedersen & nøhr-hansen 2014). most of the upper cretaceous is represented except for hiatuses in the middle turonian, lower cam panian and middle maastrichtian (figs 7, 9). palaeogene the thickness of paleocene sediments on the labrador margin varies considerably both between and within basins. this is shown by successions in the gjoa g-37, hekja o-71 and ralegh n-18 wells, all of which are in that part of the saglek basin that underlies the davis strait (figs 1, 2). within these three wells, the most complete paleocene sequence is in gjoa g-37, which has minimal danian, partial selandian and thanetian rocks. hekja o-71 appears to have selandian and tha netian strata, whereas ralegh n-18 has only thanetian rocks. wells farther south in the saglek basin contain minimal paleocene (e.g. skolp e-07, figs 1, 2, 8) or a condensed selandian and thanetian succession (e.g. gilbert f-53). in the hopedale basin, the paleocene succession is thin and discontinuous in bjarni o-82 and ogmund e-72. in north leif i-05 and ogmund e-72, only danian and thanetian strata are present. and in south labrador n-79, the selandian is con densed (fig. 8), perhaps reflecting slower sedimentation rates due to uplift. a major danian hiatus occurs in wells off west green land, with upper paleocene strata directly overlying lower bulletin37.qxp_bulletin 37 26/01/17 13.59 side 32 33 campanian rocks in ikermiut-1 and qulleq-1 (fig. 8). a thin danian section occurs, however, in nukik-2 (nøhrhansen 2003). upper selandian and thanetian strata occur in the lowermost part of kangâmiut-1 and helle fisk-1, whereas thanetian seems to be the oldest strata in nukik-1 (fig. 8). we consider this hiatus on the two mar gins to equate with the bylot unconformity of mcwhae et al. (1980) and with the base-tertiary unconformity of sinclair (1988), although the gap is greater on the green land margin. this is considered to represent the break-up unconformity marking the onset of drifting in the la b rador sea; this conforms with the conclusion of chal mers & pulvertaft (2001) that the drift phase started during the selandian. onshore in the nuussuaq basin, a lower dan ian to possibly lower selandian clastic succession is present (nøhr-hansen et al. 2002; dam et al. 2009; pedersen & nøhr-hansen 2014; fig. 3); this is overlain by a thick pile of volcanic rocks (larsen et al. 2016; fig. 3). a less obvious hiatus, involving the absence of part or all of the selandian or a condensed sequence, occurs on the labrador margin in several wells, including gjoa g-37 (as noted above) in the saglek basin and bjarni 082, north leif l-05, snorri j-90 and south labrador n-79 in the hopedale basin (figs 1, 2, 8). on the west greenland margin, danian sediments in the nukik-2 well are overlain by upper selandian sediments. at first glance, the selandian hiatus appears to be part of the bylot unconformity, but the occurrence of danian se di ments suggests a more complex scenario. in a number of wells on both sides of the labrador sea, parts of the middle eocene are marked by a hiatus or a condensed section representing the mid-eocene unconformity of dalhoff et al. (2003; fig. 8). these wells are: bjarni o-82, north leif i-05, ogmund e-72 and south labrador n-79 in the hopedale basin; gilbert f-53, gjoa g-37 and ralegh n-18 in the saglek basin; and hellefisk-1, ikermiut-1 and kangâmiut-1 off west greenland (nøhr-hansen 2003). hiatuses in the younger palaeogene and neogene are difficult to determine, largely because sedimentation is dominated by coarse clastic rocks with a low preservation potential for palynomorphs. a hiatus is recognised within the oligocene, however, on both sides of the labrador sea and it extends into the miocene in several wells. on the western margin, this hiatus seems to pri marily represent the chattian, though it may include part of the rupelian. in the southern hopedale basin, chattian strata are absent in south labrador n-79. in bjarni o-82, however, the thick rupelian interval is overlain by strata that may be of chattian age. in snorri j-90 in the northern hopedale basin, rupelian sediments appear to be directly overlain by miocene sediments (fig. 8; appendix 3.5). the rupelian is also thick in some of the saglek basin wells, such as gilbert f-53 to the south, and sediments of this age are present in gjoa g-37 and hekja o-71 to the north. in hekja o71, the thick rupelian section is overlain by up to 30 m of chattian sediment, an interval not usually pre served in the region. the oligocene hiatus of the western labrador sea pro bably corresponds to the regional unconformity pla ced at the top of the kenamu formation by balkwill et al. (1990), the baffin bay unconformity (fig. 6, in pocket, fig. 8). these authors pointed out that most of the inner and central parts of the northern labrador and south-eastern baffin margins were emergent in the middle to late oligocene, so that the oldest post-drift rocks are represented by the lower member of the mokami formation (fig. 3). off west greenland, rolle (1985) recorded oligo cene strata in hellefisk-1 and possible oligocene strata in four wells: ikermiut-1, kangâmiut-1, nukik-1 and nu kik-2 (figs 1–3, 8). however, based on more recent bio stratigraphic data, it is most likely that the oligocene unconformity extends across to the greenland side, since piasecki (2003) recorded an ypresian–serravallian hiatus in qulleq-1. the youngest palaeogene interval found by nøhr-hansen (2003) in hellefisk-1 is of priabonian age (figs 3, 8; appendix 3.19). this oligocene unconformity broadly correlates with the cessation of sea-floor spreading in the rupelian in the northern labrador sea and in the davis strait, and probably also in baffin bay. srivastava (1978) and dickie et al. (2011) considered that the cessation of seafloor spreading occurred during the priabonian (ano maly 13). however, a more likely cause for the unconformity was the development of the antarctic cir cum polar current between 33 and 30 ma, which marked the onset of antarctic glaciation. this triggered a drop in sea level that was worldwide (haq et al. 1987; haq & alqahtani 2005). neogene unconformities spanning the palaeogene–neogene bound ary and within the neogene in the labrador–baf fin seaway are difficult to identify and correlate. as noted earlier, mcwhae et al. (1980) placed the beau fort unconformity between the mokami formation and the overlying saglek formation (fig. 3); grant (1980) postulated a regional upper miocene unconforbulletin37.qxp_bulletin 37 26/01/17 13.59 side 33 34 0 ma 5 10 15 20 25 30 35 40 45 50 55 60 65 70 75 80 85 90 95 100 105 110 115 120 125 late cretaceous paleocene eocene oligocene miocene pliocene quaternary p er io d / e p o ch a ge ( st ag e) fo rm at io n m em b er r eg io n al (d ic ki e et a l. 2 0 1 1 ) l ab ra d o r m ar gi n (m cw h ae 1 9 8 1 ) early cretaceous aptian albian cenomanian turonian coniacian santonian campanian maastrichtian danian selandian thanetian ypresian lutetian bartonian priabonian rupelian chattian aquitanian burdigalian langhian serravallian tortonian messinian zanclean piacenzian mokami kenamu cartwright markland bjarni saglek upper lower upper leif lower lower upper + middle gudrid lower gudrid upper freydis lower freydis upper sn o rr i j9 0 h e k ja o -7 1 n o rt h l e if i0 5 b e au fo rt b af fin b ay b yl o t a va lo n l ab ra d o r r o b e rv al k -9 2 b ja rn i o -8 2 o gm u n d e -7 2 s. l ab ra d o r n -7 9 sk o lp e -0 7 k ar ls e fn i a -1 3 g ilb e rt f -5 3 r u t h -1 1 r al e gh n -1 8 g jo a g -3 7 s n chronostratigraphy hiatuses labrador margin preserved stratigraphic interval labrador margin lithostratigraphy bulletin37.qxp_bulletin 37 26/01/17 13.59 side 34 35 ma h e lle fis k -1 n u k ik -1 n u k ik -2 k an gâ m iu t1 ik e rm iu t1 q u lle q -1 g r o # 3 & s u rf ac e se ct io n s n . n u u ss u aq u m iiv ik -1 , sv ar te n h u k h al vø & n . n u u ss u aq , su rf ac e s e ct io n s s n late cretaceous paleocene eocene oligocene miocene pliocene quaternary p er io d / e p o ch a ge ( st ag e) early cretaceous aptian albian cenomanian turonian coniacian santonian campanian maastrichtian danian selandian thanetian ypresian lutetian bartonian priabonian rupelian chattian aquitanian burdigalian langhian serravallian tortonian messinian zanclean 0 5 10 15 20 25 30 35 40 45 50 55 60 65 70 75 80 85 90 95 100 105 110 115 120 125 n o t p en et ra te d o ff sh o re chronostratigraphy west greenland margin preserved stratigraphic interval stratigraphic record labrador 0 % % 100 100 west greenland (offshore) 0 piacenzian fig. 9. schematic representation of the preserved stratigraphic record on the labrador and greenland margins, comparing the preserved record in the sections individually and collectively with the hiatuses linked to unconformities reported in the literature. dickie et al. (2011) recognised a number of unconformities in two or more of three regions – southwest greenland shelf, the labrador shelf and the jeanne d’arc basin, offshore newfoundland (red, all three regions; green, two regions). mcwhae (1981) named five unconformities on the labrador margin; the hiatuses associated with these unconformities indicated here are estimates by mcwhae (1981) based on an unspecified timescale and thus should be considered approximate. the ‘cumulative preservation’ plots (right) depict sche ma tically the proportion of wells that contain sediments of a specified age, based on one millionyear slots; note that these calcu lations only include wells that extend down to/beyond the relevant stratigraphic level. bulletin37.qxp_bulletin 37 26/01/17 13.59 side 35 36 mity. in the wells we have examined, a middle to late miocene hiatus occurs in the hopedale basin. for ex ample in bjarni o-82 (fig. 8), the serravallian is overlain by pliocene–pleistocene sediments, indicating that all of the upper miocene is absent. a similar situation seems to exist in south labrador n-79, where plio cene–pleistocene sediments apparently overlie the mid dle miocene succession (fig. 8). furthermore, in snorri j-90 in the northern part of the hopedale basin, the upper middle miocene is missing (fig. 8). evidence in the saglek basin is more tenuous. in gilbert f-53, the sparse palynomorph assemblage pro vides inconclusive evidence, but the thin nature of the miocene strongly suggests a significant hiatus. un for tunately, in gjoa g-37 and hekja o-71, samples were not available from the upper drilled interval and the uppermost samples analysed are of rupelian and early miocene age, respectively. in offshore west greenland, piasecki (2003) recorded middle miocene as the oldest neogene interval in the kangâmiut-1 well. a similar situation may exist in the nukik-1, nukik-2, ikermiut-1 and hellefisk-1 wells, although no marker species were recorded among the sparse palynomorph assemblages in the samples from the upper part of the ikermiut-1 and hellefisk-1 wells. in summation, there is convincing evidence that some of the eocene and all of the oligocene and lower miocene are missing on the west greenland margin, so that middle miocene sediments directly overlie those of ypresian to priabonian age. on the western margin of the seaway, however, there is a middle to late mio cene hiatus, especially in the hopedale basin (fig. 8). this hiatus may be equivalent to the beaufort un con formity (fig. 3) of mcwhae et al. (1980) and mcwhae (1981). palaeoenvironmental results general considerations the general methodology of palaeoenvironmental in ter pretation in this study is reviewed in this and the following sections. the initial approach for some wells was to develop palaeoenvironmental plots showing the ratio of mio spores to dinocysts, having separated out acritarchs and other organic-walled organisms, such as massullae of azolla. this ratio provides a rough estimate of distance from the shoreline, although it can be misleading be cause of the dominance of bisaccate pol len far offshore and the general decrease in dinocysts in oligotrophic zones (see below). if there are no dinocysts present it was assumed, perhaps incorrectly, that the sediments were non-marine. subsequently, the focus of the pa laeo ecological study switched to fluctuations in dino cyst assemblages and how the palaeoenvironmental de terminations compared with previous results in the labrador–baffin seaway region, the palaeoenvironmental interpretations are based on detailed counts for six wells: bjarni o-82, gilbert f-53, gjoa g-37, hejka o-71, snorri j-90 and south lab rador n-79, but are augmented by general observations from other wells. for these six wells, counts of 100 to 200 specimens per sample were made where possible. dino cyst taxa considered significant are listed in table 1, and the plots in appendix 2 include a column highlighting palaeoenvironmental interpretations. the co lumn has five subdivisions: non-marine, marginal ma rine, in ner neritic, outer neritic and open ocean, and the environmental assessment through time is indicated for each well by a curve in this column. both quantitative and qualitative data have been assessed in developing the palaeoenvironmental curves. for example, the presence of two or three specimens of impagidinium was deemed sufficient to interpret the palaeoenvironment as open ocean (dale 1996), although as zonneveld et al. (2013) noted, impa g idinium is not invariably restricted to open-ocean environments in recent sediments. however, to strengthen the interpretations, occurrences of impagidi nium in the labrador–baffin seaway wells were com pared with plank tonic foraminiferal data. palaeoenvironmental curves for the palaeogene successions in greenland wells (hellefisk-1, ikermiut1, kangâmiut-1, nukik-1 and nukik-2) published in rasmussen et al. (2003) are included in the present study (appendices 3.15–3.19). the curves indicate the presence of littoral/lagoonal (transitional), inner neritbulletin37.qxp_bulletin 37 26/01/17 13.59 side 36 37 ic, middle neritic, outer neritic and upper bathyal pa laeo environments, and were interpreted from the pa ly no morph content (mainly dinocysts and their species rich ness, freshwater algae, fungal spores, miospores, plant tissues) in association with data from other fossils (mainly foraminifera, radiolarians, diatoms, ostracods, gas tro pods, bivalves and calcareous nannofossils). in gener al, the pa laeo environmental signals from different fossil groups show similar trends (rasmussen et al. 2003); the discrepancies that do occur may be due to the fact that most of the samples studied are cuttings. use of dinocysts for palaeoenvironmental interpretations can be misleading. dinocyst assemblages tend to reflect distance from shore rather than water depth, which is why ‘open ocean’ is used in appendix 2 rather than ‘bathyal’. since the relationship between water depth and distance from shore is not a simple one – varying for example in relation to shelf width – the curves presented here should be treated with some circumspection. further caution is needed because, as with the biostratigraphy, the palaeoenvironmental in ter pretations are based on data from cuttings that can lead to misinterpretation due to downwell contamination. however, the curves, most of which have been reported previously (williams 2007a–e), do show gen er al trends and agree well with interpretations from, for example, foraminiferal and sedimentological data (bu jak davies group 1987; miller & d’eon 1987). palaeoenvironments and dinocysts: previous studies pioneering studies in interpreting palaeoenvironments from dinocyst assemblages were by gocht (1969) and downie et al. (1971). one early approach involved the ‘gonyaulacacean ratio’, based on an innovative study of the campanian bearpaw formation of alberta by har land (1973). this author defined the ratio as the num ber of species that have a gonyaulacacean affinity di vided by the number of species having a peridiniacean affinity. thus the ratio reflects number of species rather than number of specimens. harland assumed that when the number of gonyaulacacean species is higher than the number of peridiniacean species, more open-marine conditions prevailed. the interpretations based on this approach were in agreement with the foraminiferal data for the bearpaw formation. in our studies, however, it was found that specimen counts were more meaningful than number of species. palaeoenvironental studies using dinocysts have ex panded significantly in the past thirty years. important works include those by köthe (1990), brink huis (1994), dale (1996), stover et al. (1996), powell et al. (1996), jaramillo & oboh-ikuenobe (1999), sluijs et al. (2005), sluijs et al. (2008), sluijs & brinkhuis (2009), guerstein et al. (2008), lebedeva (2010) and schreck et al. (2012). the comprehensive atlas of the occurrences of modern cyst-forming dinoflagellates by zonneveld et al. (2013) is also very helpful in projecting modern distri butions into the past. from a review of the literature, it is clear that some palaeoenvironments are more clearly interpreted than others from dinocyst assemblages, and some taxa give clearer signals than others. one such signal comes from the late early cretaceous ceratiacean cysts with adnate apical archaeopyles – nyktericysta and vesperopsis. mac rae (1996) noted that several studies – including wight man et al. (1987), bint (1986), banerjee & davies (1988) and leckie & singh (1991) – had suggested that at least some species of nyktericysta and vesperopsis were specialised for brackish or freshwater environments. nøhr-hansen (1992, 2008), zippi (1998), dolby et al. table 1. palaeoenvironmental preferences of dinocyst taxa in the labrador–baffin seaway coastal – marginal marine inner neritic outer neritic open ocean cannosphaeropsis impagidinium nematosphaeropsis pterodinium cerodinium cleistosphaeridium cordosphaeridium hystrichokolpoma hystrichosphaeridium operculodinium phelodinium spiniferites areoligera cleistosphaeridium cribroperidinium deflandrea dinogymnium glaphyrocysta heterosphaeridium micrhystridium* phthanoperidinium wetzeliella eocladopyxis heteraulacacysta homotryblium micrhystridium* nyktericysta polysphaeridium tuberculodinium vesperopsis * acritarch bulletin37.qxp_bulletin 37 26/01/17 13.59 side 37 38 (2013) and pedersen & nøhr-hansen (2014) similarly concluded that vesperopsis mayi and nyktericysta (as bal mula) tripenta and similar ceratiacean taxa seem to be associated with non-marine to marginal marine en vironments. in the cenozoic, some of the most useful dinocyst taxa for indicating marginal marine palaeoenvironments are members of the family goniodomaceae, in cluding homotryblium tenuispinosum (brinkhuis 1994) and eocladopyxis (sluijs et al. 2005). crouch et al. (2001, 2003) and sluijs et al. (2005) also noted that abundances of the peridiniacean (wetzelielloidean) apec to dinium spp. were common in marginal marine settings in the late paleocene – early eocene. brinkhuis (1994) con sidered areoligera, glaphyrocysta, homotryblium, oper cu lo dinium, spiniferites and areosphaeridium to define innershelf palaeoenvironments in the early cenozoic. an example of conflicting palaeoenvironmental in ter pretations involves the findings of heilmannclausen (1994) and powell et al. (1996). heil mannclausen (1994) considered areoligera gippingensis to be indicative of offshore marine enironments. he noted (heilmann-clausen 1994, p. 53) that areoligera gippingensis is abundant throughout the latest selandian – early thanetian alisocysta margarita zone in denmark, where “the entire zone occurs in a clearly offshore set ting”. powell et al. (1996), however, described an areo ligera-dominated assemblage (probably equivalent to the areoligera gippingensis complex of this study) from the thanetian type section of southern england. in their pegwell bay section, they noted three intervals where areoligera was ‘superabundant’ and three where it was ‘abundant’. according to powell et al. (1996), such horizons denote restricted high-energy, marginal marine settings typical of a transgressive regime. they believed that the richest samples were close to the most condensed interval or maximum flooding surface. as discussed further below, it appears that assemblages rich in specimens of the areoligera gippingensis complex can give mixed messages. downie et al. (1971) postulated that their ypresian wetzeliella association was estuarine. this association was dominated by species of wetzeliella and (perhaps) deflandrea. dominant taxa included apectodinium (as wetzeliella) homomorphum, apectodinium (as wetze liel la) parvum and deflandrea phosphoritica. firth (1996), who studied high-latitude north atlantic dinocyst assemblages from odp hole 913b, found relative abundance peaks of deflandrea spp., together with phthanoperidinium spp., in strata that are primarily diatomand radiolarian-rich biosiliceous oozes that clearly originated in a deep-water, open-ocean palaeo environment. firth (1996) concluded that these peaks could reflect high palaeoproductivity events. the interpretations of downie et al. (1971) and firth (1996) demonstrate that similar dinocyst assemblages may re flect conditions such as nutrient enrichment (usually resulting from upwelling) that may occur in a range of marine settings, as is the case today. more clarity was provided by the studies of brink huis et al. (2003) and sluijs et al. (2003), who showed that lithological, geochemical, grain-size and diatom data indicate that high abundances of deflandrea and phthanoperidinium denoted shallow-water palaeoenvironments. köthe (1990) also concluded that common to abundant occurrences of phthanoperidinium suggest an inner neritic palaeoenvironment. clearly, the palaeo environmental interpretation of dinocyst assemblages is best unravelled with support from other evidence, such as lithological and benthic foraminiferal data. harding (1990), eshet et al. (1994), brinkhuis et al. (1998) and van mourik et al. (2001) all suggested that high abundances of palaeoperidinium cysts indicate higher nutrient levels that reflect terrigenous input and therefore palaeoenvironments probably nearer to shore. however, since nutrient enrichment can be related to upwelling processes, as shown later for shelf-edge settings, more than one interpretation is possible for the abundant occurrences of palaeoperidinium. oligotrophic palaeoenvironments are commonly associated with mid-shelf settings, where nutrients are low, relative to more nutrient-rich proximal settings, and distal locations where there is upwelling (sluijs et al. 2005). in the labrador margin wells, oligotrophic palaeoenvironments can usually be determined from the dramatic drop in numbers of dinocyst specimens, which is commonly accompanied by a marked relative increase in the number of miospores. this is demonstrable in the interval 2020–1450 m in gilbert f-53 (appendix 2.4), in which dinocyst specimen counts are extremely low (williams 2007e) and where the fora miniferal data confirm that the palaeoenvironment was middle neritic (miller & helenes 1989c). for example there are nine dinocysts compared to 95 miospores at 1940–1930 m. it is extremely difficult to determine which dinocyst species are characteristic of outer neritic palaeoenvironments. in their study of early eocene assemblages, downie et al. (1971) postulated that their hystricho sphaeridium association is indicative of open-sea pa laeoenvironments. this association included species of achomosphaera, cordosphaeridium, hystricho sphaeri dium bulletin37.qxp_bulletin 37 26/01/17 13.59 side 38 39 and spiniferites. brinkhuis (1994) came to similar but more detailed conclusions: in his priabonian outer neri tic dinocyst assemblages, he included the taxa cleisto sphaeridium, spiniferites, areosphaeridium, oper culo dinium, nematosphaeropsis and cannosphaeropsis. other taxa indicating outer neritic palaeoenvironments are species of hystrichokolpoma; this genus, which can also be associated with impagidinium, is suggestive of openocean conditions (williams 2007e). support for this interpretation is provided by foraminiferal data for the gilbert f-53 well (bujak davies group 1987). open-ocean palaeoenvironments are characterised by the occurrence of impagidinium (sluijs et al. 2005). this conclusion is based on the pioneering study of wall et al. (1977). fortunately impagidinium occurs in the late cretaceous, at least in the maastrichtian in this study, as well as ranging throughout the cenozoic. hystrichokolpoma, as discussed above, can also indicate outer neritic to open-ocean palaeoenvironments, just as areoligera (herein areoligera gippingensis complex) abun dances seem to do in the paleocene. this is confirmed by foraminiferal data in the labrador margin wells (miller & helenes 1989b; williams 2007e). previous palaeoenvironmental interpretations from the labrador– baffin seaway published palaeoenvironmental conclusions on the western labrador sea wells are based on foraminifera (gradstein & williams 1976; gradstein & srivastava 1980; bujak davies group 1987; miller & helenes 1989a–c; ainsworth et al. 2014), palynomorphs (grad stein & williams 1976, williams 2007a–f) and nannofossils (crux & gard 2004). similar findings for west greenland utilise foraminiferal, nannofossil and palynological studies of the lower palaeogene (rasmussen et al. 2003; sheldon 2003) and dinocyst studies of the neogene (piasecki 2003). using foramineral and palynological data from wells, gradstein & williams (1976) concluded that depositional environments on the labrador margin evolved from non-marine in the early cretaceous to possibly neritic in the late cretaceous – paleocene, then to bathyal in the eocene, neritic in the oligocene–mio cene, and littoral and non-marine in the pliocene–plei stocene. an apparent conflict is evident between the palynological and foraminiferal data from the eocene, where the foraminifera indicate bathyal palaeoenvironments and the dinocysts suggest inner neritic conditions; the most plausible explanation seems to be that the dinocysts were redeposited. in a reappraisal of the data, gradstein & srivastava (1980) defined four depositional palaeoenvironments for labrador margin wells, again based on foraminifera and palynomorphs. these were: non-marine; shallow neritic (taken to include marginal marine to inner shelf, with water depths less than 100 m); deep neritic (with water depths ranging from 100–200 m); and bathyal or upper slope (with water depths from 200–1000 m). according to these authors, marine sedimentation start ed in the middle late cretaceous in the labrador sea, but the main transgression began during the maa strichtian, thus coinciding with the onset of sea-floor spreading during anomaly 32. neritic to bathyal con ditions persisted throughout paleocene and eocene times, with the deepest water (bathyal) palaeoenvironments in the eocene reflecting rapid subsidence at that time. near the eocene–oligocene boundary, at the end of sea-floor spreading, palaeoenvironments changed from deep to shallow marine, the latter conditions per sisting through the later cenozoic. according to grad stein & srivastava (1980), the only exception was in kangâmiut-1 where deeper water environments per sisted into the neogene. foraminiferal analyses of cores from leg 105, site 645 in baffin bay (srivastava et. al. 1987) showed that an upper slope palaeoenvironment persisted there in the miocene. the dinocysts from the same interval indicated open-marine conditions. higher in the miocene, open-marine and neritic dinocyst species occur, probably reflecting redeposition of the latter. at site 646, palaeoenvironments were generally lower bathyal. results and interpretations cretaceous the observations made during this study support earlier findings that non-marine conditions prevailed du r ing the early cretaceous on what is today the labrador margin (gradstein & williams 1976); this interpretation is based on the fact that many assemblages in this part of the section consist exclusively of miospores. balkwill et al. (1990) considered the bjarni formation, which they dated as early barremian to albian, to represent lacustrine deposits. however, the occasional occurrence of species of nyktericysta and vesperopsis on both sides of the seaway suggests a non-marine to marginal marine setting, as does the presence of subtilisphaera bulletin37.qxp_bulletin 37 26/01/17 13.59 side 39 40 spp. in bjarni o-82 and pseudoceratium sp. in roberval k-92. the interval containing these dinocysts was interpreted as non-marine by miller & d’eon (1987), but williams (2007a) interpreted it in part as marginal marine, based on the few dinocysts present. the overlying rocks contain other aptian dinocyst species, especially tenua hystrix, which is known only from shallow marine palaeoenvironments. thus, marine environments extended back to about 120 ma. possible evidence for earlier marine conditions in the davis strait area are the reworked late jurassic dinocysts peris seia sphaeridium pannosum and fromea tornatilis recorded from qulleq-1 by nøhr-hansen et al. (2000). nyktericysta is present and sometimes common in the aptian of south labrador n-79, the albian–ce nomanian in ogmund e-72, and north leif i-05 in the hopedale basin (nøhr-hansen in sønderholm et al. 2003b; nøhr-hansen 2004b; williams 2007b) and in samples from nuussuaq and disko in the nuussuaq basin of onshore west greenland (nøhr-hansen 2008; pedersen & nøhr-hansen 2014). these onshore green land occurrences accord well with labrador sea observations, and suggest that conditions in the entire region fluctuated between shallow marine and lagoonal to la custrine in the aptian?/albian to cenomanian. inner neritic conditions became widespread in the turo nian/coniacian to santonian interval. dinocysts are usually abundant in such settings, especially in the vicinity of discharge from rivers, where nutrients are plentiful. taxa that appear to indicate inner neritic palaeoenvironments in this interval include hetero sphaeridium difficile, circulodinium distinctum, chatan giella and gillinia hymenophora. these conclusions are based on the studies of shallow-water assemblages from bylot island (g.l.williams, unpublished data), in which foraminifera are absent. more open-ocean, presumably deeper water palaeo environments developed in the campanian and maa strichtian, as shown in gilbert f-53 in the saglek basin and bjarni o-82 and south labrador n-79 in the hopedale basin. foraminiferal data for all three wells indicate that the palaeoenvironments were bathyal (bu jak davies group 1987). cenozoic intervals with occurrences of phelodinium kozlowskii tend to also contain foraminifera that indicate a bathyal palaeoenvironment (bujak davies group 1987); hence, the former may indicate an outer neritic or open-ocean palaeoenvironment. phelodinium kozlow skii occurs in some samples between 3200 and 2740 m in gilbert f53, an interval that is of maastrichtian (3250–3120 m) to danian (3100–2740 m) age (wil liams 2007e). this deeper water interpretation seems to be in conflict with the presence of the hiatus between the cretaceous and cenozoic sediments in some of the wells. however, upper maastrichtian and danian to selandian strata, including transitional cretaceous–ce nozoic strata, oc cur in nuussuaq. other common dinocyst taxa in the paleocene are alterbidinium spp., areoligera gippingensis, cerodinium diebelii, cerodinium speciosum, glaphyrocysta divaricata, hystrichosphaeridium tubiferum, palaeocystodinium golzowense, palaeoperidinium pyrophorum and trithyro dinium evittii. the pollen pinuspollenites is also extreme ly abundant in the paleocene. that palaeoperidinium pyrophorum is considered to indicate outer neritic to open-ocean palaeoenvironments is based on foramini feral data (e.g. gilbert f-53). this interpretation seems at odds, however, with the conclusions of several au thors (e.g. eshet et al. 1994; brinkhuis et al. 1998; van mourik et al. 2001), who postulated that high abundances of palaeoperidinium (peridinioid) cysts in dicate higher nutrient levels, thus reflecting terrigenous input and a closer proximity to shore. dale & fjellså (1994) pointed out that the assumption that peridinioid cysts are invariably heterotrophic is not always warranted and can lead to erroneous interpretations; they also noted that modern heterotrophic dino flagellates are not re stricted to areas of high pro ductivity. in addition, nu trient-rich conditions can develop in various, contrasting settings including inshore areas with abundant terrigenous input and shelf-edge settings with upwelling, so that different interpretations are en tirely possible. the gilbert f-53 well (3200–2770 m) provides an example in which an abundance of palaeoperidinium pyrophorum is related to an offshore setting rather than to a coastal environment. there is a marked increase in the abundance of this species from 3260 to 2740 m (excluding the interval 2860–2760 m) indicative of increasing trophic resources. over the same interval, the gonyaulacacean ratio, which is considered to be an indicator of distance from shore, varies from 0.5 to 1.4, suggesting progressively more open-marine conditions. the foraminiferal data from this well indicate that the interval from 3200 to 2740 m represents a bathyal setting (bujak davies group 1987). miller & helenes (1989b) interpreted the interval from 3330 to 3230 m as middle to outer neritic, and from 3230 to 2520 m as bathyal, with the interval from 2940 to 2750 m as bulletin37.qxp_bulletin 37 26/01/17 13.59 side 40 41 possibly outer neritic. similarly, the dinocyst taxa re corded from about 3250 to 2750 m indicate outer-shelf palaeoenvironments; the presence of impagidi nium in some samples suggests open-ocean conditions (dale 1996). during the danian, the palaeoenvironmental setting of south labrador n-79 was presumably similar to that of gilbert f-53; the former well also shows high abundances of palaeoperidinium pyrophorum in association with impagidinium. thus the abundance of palaeo peridinium pyrophorum in both wells is best explained by high nutrient levels at the shelf edge related to deepwater upwelling, conditions that are ideal for plankton whether autotrophic or heterotrophic. several dinocyst groups proliferated during the pa laeogene, an example being the areoligeraceans, especially the genera areoligera and glaphyrocysta. species of these two genera range through parts of the late cre taceous and palaeogene, and the species areoligera gip pingensis and glaphyrocysta divaricata are especially pro minent in the paleocene. the two species are hard to distinguish – we would restrict areoligera gippingensis to specimens with basal ridges at least partially connecting processes in process complexes and glaphyrocysta divaricata to specimens lacking such ridges (see fen some et al. 2016). most records in the literature understandably just record the complex as areoligera gip pingensis or areoligera spp. (although our experience on the canadian margin is that most specimens belong to glaphyrocysta divaricata). as noted above, we refer to the two species collectively herein as the areoligera gip pingensis complex. specimens of the areoligera gippingensis complex are among the most common dinocysts in the paleocene of the labrador margin. these forms are abundant in the danian–selandian of south labrador n-79 in the hopedale basin, in the maastrichtian (where they are interpreted as caved) to thanetian of gilbert f-53, and in the selandian–thanetian of gjoa g-37 and the thanetian of karlsefni a-13 in the saglek basin. in these wells, peak abundances of the complex can occur throughout the paleocene. similar peaks of this species complex have been observed elsewhere, with contrasting interpretations, for example on the scotian margin (fensome et al. 2008), in denmark (heilmann-clausen 1994) and in the type thanetian of southern england (powell et al. 1996). members of the areoligera gippingensis complex are gonyaulacaceans and thus more likely to be the cysts of autotrophic dinoflagellates. their presence in dominant numbers in wells of the seaway suggests that this was in an area of high nutrient concentrations, presumably but not necessarily related to upwelling. if the complex gravitated towards regions of high productivity, it could have occupied a coastal zone or been near to the shelf margin, as discussed earlier in relation to pa laeo peridinium pyrophorum. independent studies of the intervals containing high abundances of the areoli gera gippingensis complex in the labrador margin wells (bjar ni o-82 and south labrador n-79) infer an outershelf to bathyal environment (bujak davies group 1987; miller & d’eon 1987). it is concluded that high abundances of areoligera and glaphyrocysta can indicate either outermost shelf or innermost shelf settings, the latter related to a transgressive regime. accordingly such abundances must be treated with care in determining palaeoenvironments, not least since their great abundances would surely have guaranteed their dispersal in currents between environments. off west greenland, the areoligera gippingensis complex is abundant in the predominantly outer neritic selandian interval in the kangâmiut-1 and ikermiut-1 wells (appendices 3.17 and 3.18 respectively), and present to common in the marginal marine to inner neritic selandian interval in hellefisk-1, nukik-1 and nukik-2 (appendices 3.19, 3.15, and 3.16 respectively; rasmussen et al. 2003). such varied environmental interpretations may also explain the occurrence patterns of the complex on the canadian side of the seaway. the interpretation of relatively high abundances of areoligeraceans has been mixed. sluijs et al. (2008) argued that peridinioid dinocysts are less sensitive indicators of proximity to shore than gonyaulacoids; within the gonyaulacoids, these authors suggested that the most reliable indicator for the paleocene–eocene interval is the s/a index (spiniferites/spiniferites + areo ligera). sluijs et al. (2008) argued that a high relative abundance of areoligera (a low s/a ratio) indicated inner neritic palaeoenvironments, whereas a high re lative abundance of spiniferites specimens (a high s/a ratio) indicated outer neritic palaeoenvironments. there are two provisos with this approach: firstly, that glaphyrocysta counts should be included with areo ligera; and secondly, that any conclusions based on this approach should, where possible, be supported by plank tonic foraminiferal data. building on the research of sluijs et al. (2008), sluijs & brinkhuis (2009) equated the dominance of the areoligera complex (presu mably including glaphyrocysta) with inner neritic, high-energy environments and the spiniferites complex with neritic deposits, with relative abundances increasing in outer neritic palaeoenvironments (see also brink huis 1994 and pross & brinkhuis 2005). from their bulletin37.qxp_bulletin 37 26/01/17 13.59 side 41 42 study of paleocene–eocene assemblages on the new jersey shelf, sluijs & brinkhuis (2009) concluded that abundant areoligera appears to be consistently related to third-order transgressive systems tracts, indicating ri sing sea level in a neritic palaeoenvironment. in debates about the relationship between dinocyst assemblages and palaeoenvironment, one factor missing from the discussion may be evolutionary contingency. the striking acmes of the areoligera gippingensis com plex in the paleocene, especially the thanetian, may re flect ideal conditions (including climate) for the pro li feration of these species at the time of their evolutionary appearance; under such conditions, the species spread widely, possibly as ecological generalists. the species were present still in the early eocene, but did not dominate assemblages as they had in the paleocene, possibly due to competition from other ‘generalists’. thus, environmental interpretations of the areoligera gippingensis complex may be specific to particular stages in the evolutionary development of the group, and may change with time. open-ocean conditions persisted into the early eo cene, with the foraminifera indicating deep water (gradstein & srivastava 1980). there are some differences between the palaeoenvironmental interpretations for the eocene based on dinocysts and foraminifera: the dinocysts indicate bathyal palaeoenvironments with some neritic excursions, changing to outer to middle neritic in the priabonian, whereas the foraminifera in dicate more consistent bathyal palaeoenvironments changing to outer neritic in the priabonian (gradstein & williams 1976). collectively, what were the palaeoenvironmental ex tremes during the paleocene–eocene in the seaway as reflected in the offshore wells? this is best illustrated by comparing the data from gjoa g-37 (appendix 2.6), hekja o-71 (appendix 2.5) and ralegh n-18 (not illu strated here). in gjoa g-37, assemblages from the paleo cene and eocene suggest predominantly outer neritic palaeoenvironments, with open-ocean interludes, pre sumably denoting deeper water palaeoenvironments. in contrast, assemblages of the same age in hekja o-71 and ralegh n-18 suggest primarily marginal marine to inner neritic palaeoenvironments, with occasional episodes of non-marine deposition. thus, the lowermost eocene suc cession in ralegh n-18 contains common specimens of the freshwater alga pediastrum (nøhr-hansen 2004a) and common taurodinium granulatum, which are as sumed to be autochthonous. taurodinium granulatum is also common in hekja o-71, together with scattered specimens of pediastrum. however, the rare specimens of these taxa in gjoa g-37 (nøhr-hansen in sønderholm et al. 2003b; nøhr-hansen 2004a) are inferred to have been transported into the depositional environment. taurodinium granulatum has previously been recorded as “gen et. sp. indet” by hjortkjær (1991) and piasecki et al. (1992) from the syn-volcanic, probably lacustrine selandian deposits at assoq and on disko, west green land, and from pre-basaltic thanetian–?ypresian lacustrine deposits from the kangerlussuaq basin in south-eastern greenland (nøhr-hansen 2012). thus, the common occurrence of freshwater algae and lacustrine to brackish-water dinocysts in hekja o-71 and ralegh n-18 suggests a position close to the paleocene– eocene palaeoshoreline. the contrast between the hekja o-71 and ralegh n-18 wells on the one hand and the gjoa g-37 well on the other, is compatible with their relative positions within the basin (figs 1, 2). the same pattern has also been recognised in offshore west greenland wells. rasmussen et al. (2003) showed that paralecaniella indentata is common in the lowermost thanetian and lowermost ypresian of hellefisk-1, nukik-1 and nukik-2. nøhr-hansen (2004a) recognised an acme of paralecaniella indentata at the same stratigraphic level in ralegh n-18. paralecaniella ap pears to be most abundant in marginal marine successions but has a tolerance for inner neritic conditions, presumably in the vicinity of river estuaries or deltas, where there is an influx of freshwater (elsik 1977; powell et al. 1996). however, lebedeva (2010) considered that paralecaniella preferred coastal, high-energy marine settings of normal salinity, with sufficient oxygen levels. an alternative explanation is that the specimens were washed in and are thus allochthonous (rasmussen et al. 2003). the common occurrence of the freshwater alga pediastrum at the same stratigraphic level in the outer neritic deposits in kangâmiut-1 and ikermiut-1 was interpreted as evidence of redeposition, probably by turbidites (rasmussen et al. 2003). an intriguing aspect of labrador sea – davis strait palynomorph assemblages is the common occurrence of azolla in some ypresian samples – especially in gjoa g-37, where it occurs with common specimens of the freshwater alga pediastrum between 1890 and 2040 m (nøhr-hansen in sønderholm et al. 2003b; appendix 3.13). the same pattern is seen in the ypresian of ralegh n-18 (appendix 3.12), but only pediastrum is common in the ypresian of hekja o-71 (appendix 3.11; nøhr-hansen 2004a). azolla is a small moss-like, freshwater to brackish-water fern famous for its nitrogen-fixing capability (van kempen et al. 2012). its modern distribution indicates that it prefers much bulletin37.qxp_bulletin 37 26/01/17 13.59 side 42 43 warmer temperatures than occur today in the la bra dor–baffin seaway and suggests the possibility of extensive freshwater lakes. massive concentrations of dead azolla accumulating at the bottom of the lakes could be preserved in sediment, and given the right combination of time and temperature, may yield significant hydrocarbon source rocks. the occurrences of azolla in la b rador sea – davis strait wells are restricted to a narrow time interval at the top of the ypresian. azolla has not been recorded in the offshore west greenland wells. brinkhuis et al. (2006) recorded numerous specimens of azolla in a core from the eastern end of the lomonosov ridge in the arctic ocean; they suggested that the ypresian records of azolla in the labrador– baffin seaway were potentially derived from natural freshwater overspill from the arctic ocean, which they proposed was periodically characterised by fresh surface waters. we consider it difficult, however, to visualise a body of water as large as an ocean with a surface that would be calm for long enough to allow azolla to become established. even if brinkhuis et al. (2006) are correct in postulating episodic fresh surface waters in the arctic during the ypresian, outlets into the atlantic ocean would probably have been via shallow seas adjacent to east greenland, not via the la b rador–baffin seaway. this observation is based on paleogeographic maps in monger et al. (2014) and fen some et al. (2014; fig. 10). azolla occurs in wells off the canadian margin but has not been found in any wells off west green land. this distribution may reflect dif ferential drainage into the seaway, such that large rivers on the canadian side introduced azolla into the seaway whereas rivers on the greenland side were too small or non-existent. the palynomorph data indicate a dramatic change in palaeoenvironments from the early to the middle and late eocene of the labrador margin and western davis strait, resulting from, or accentuated by, either a significant drop in sea level or uplift. this ultimately led to inner neritic to marginal marine conditions in the rupelian. the changes may be related to one or more of three events: the end of the drifting phase (within anomaly 13, approximately at the priabonian–rupelian boundary), the development of the antarctic ice sheet (zachos et al. 2008), or global (eustatic) sea-level fall in the oligocene (haq et al. 1977, 1987). throughout the rest of the oligocene and during the miocene to early pliocene interval, the most useful indicator of shifting palaeo environments is the miospore/dinocyst ratio. this and other evidence indicate that inner neritic to marginal marine palaeoenvironments prevailed in the vicinity of several wells: bjarni o-82 (appendix 2.1), snorri j-90 (appendix 2.3), and south labrador n-79 (appendix 2.2) in the hopedale basin; gilbert f-53 (appendix 2.4), and hekja o-71(appendix 2.5) in the saglek basin (figs 1, 2). in offshore west greenland, palynomorphs from the qulleq-1 well indicate that the middle miocene succession was deposited in an open-marine palaeoenvironment represented by the most diverse dinocyst assemblages of the entire neogene; a similar setting is indicated for the upper miocene to lower pliocene succession based on the common presence of impagidinium spp. (piasecki 2003). the peak in dinocyst species richness and abundance corresponds to the end of the middle miocene climatic optimum (zachos et al. 2001). n o r t h a m e r i c a 30°n 30°n laurasia n 85 ma 60 ma 35 ma nn 30°n n o r t h a m e r i c a fig. 10. palaeogeography of north america for the late cretaceous (85 ma), paleocene (60 ma) and eocene (35 ma), showing the evolution of the labrador–baffin seaway during that timespan. according to these reconstructions, the seaway was not directly connected with the arctic basin during the early cenozoic, which has a major bearing on the source of the azolla specimens found in some wells (see text for discussion). from monger et al. (2014) and fensome et al. (2014); maps courtesy of ron blakey. bulletin37.qxp_bulletin 37 26/01/17 13.59 side 43 44 palaeogeography, palaeoclimatology and palaeoceanography during the cretaceous and palaeogene, the labrador sea region was considerably warmer than it was during the neogene. evidence is sparse for the cretaceous, with exception of the maastrichtian. the initial findings were by gradstein & srivastava (1980), who recognised two principal influxes of warm-temperate plank tonic fora minifera – during the maastrichtian and in the early– middle eocene. later research drew attention to the several transient global warming or hyperthermal phases in the late paleocene and early eocene, especially the paleocene–eocene thermal maximum (petm; ken nett & stott 1991; koch et al. 1992). zachos et al. (2008) highlighted several other warm phases, including the two-million-year-long early eocene climatic op timum at 53–51 ma, the shorter mid-eocene climatic optimum at around 41 ma, and the fleeting eocene ther mal maximum 2 at around 52.5 ma. cooler epi sodes also occurred in the eocene, especially towards the end of the epoch, with the onset of antarctic glaciation at 35 ma. in the neogene, a moderate climatic optimum occurred in the middle miocene. early cretaceous dinocyst assemblages in the la b rador sea region are too sparse to give any clues to the climatic conditions existing there at that time. the more northerly early late cretaceous assemblages show some differences from those on the scotian margin and grand banks; most notable is the presence of hetero sphaeridium difficile, which is common in the conia cian section in bjarni o-82 and south labrador n-79 in the hopedale basin, in shallow cored boreholes from western baffin bay (maclean et al. 2014), and in upper turonian – lower coniacian strata from the umiivik-1 borehole on svartenhuk halvø (dam et al.1998; figs 1, 2). the available records (e.g. nøhr-hansen 1996; dam et al.1998; pedersen & nøhr-hansen 2014; rad macher et al. 2014) and unpublished data (g.l. wil liams) from upper cretaceous rocks of bylot island and baffin bay suggest that heterosphaeridium difficile is a high-latitude species. the assemblages also contain many of the species described by manum & cookson (1964) from graham island, arctic canada, who also recorded heterosphaeridium difficile. similarities with other dinocyst assemblages can be determined from lentin & williams (1980), who de scribed provincialism in campanian peridinialean dino cysts. these authors defined three assemblages, which they related to climatic belts: the malloy or tropical– subtropical suite, the williams or warm-temperate suite and the mcintyre or boreal suite. common com ponents of the malloy suite at generic level include andalusiella, cerodinium, lejeunecysta and sene gali nium; the williams suite is characterised by alter bi dinium, isabelidinium, spinidinium, trithyrodinium, and smaller species of chatangiella. diagnostic taxa of the mcintyre suite are laciniadinium and the larger taxa of chatangiella. the mcintyre suite is present in arctic canada, the mackenzie delta, saskatchewan, alberta, south dakota, wyoming, the northern north sea (costa & davey 1992) and west greenland (nøhrhansen 1996). expanding upon the concepts of lentin & williams (1980), mao & mohr (1992) defined a helby suite, which included diagnostic dinocyst taxa for the campanian–maastrichtian interval in the higher latitudes of the southern hemisphere. the helby suite is characterised by high abundances of the peridinioid dinocyst genera isabelidinium, chatangiella and nelsoniella. as in the mcintyre suite, specimens of cha tan giella are large and there are few specimens of an dalusiella and senegalinium. campanian assemblages on the labrador margin seem to fall into the williams suite, although large cha tan giella specimens (characteristic of the mcintyre suite) occasionally occur. dinocyst assemblages from the bu chan gulf (fig.1) on the north-eastern baffin mar gin show mcintyre-suite affinities (maclean & williams 1980; balkwill et al. 1990; maclean et al. 2014), but with some elements of the williams suite. these occurrences indicate that the climate in baffin bay during some periods of the late cretaceous was warm temperate. according to gradstein & srivastava (1980), the cam panian–maastrichtian planktonic foraminifera from nuussuaq show affinities with coeval assemblages from the north atlantic; this is based on the common occurrence of some poorly preserved taxa. these authors also noted, however, that late cretaceous climates were more equable than those of today and that species were cosmopolitan, with no specific high-latitude taxa. ceno manian to maastrichtian ammonites have affinities with coeval taxa from the western interior seaway of north america and the atlantic (birkelund 1965; williams & stelck 1975), thus partially paralleling the distributions shown by the dinocysts. bulletin37.qxp_bulletin 37 26/01/17 13.59 side 44 45 maastrichtian strata of the labrador margin are characterised by the presence of palynodinium grallator and isabelidinium cretaceum. palynodinium grallator is a ubiquitous species but isabelidinium cretaceum is most abundant at higher latitudes (askin 1988; bowman et al. 2012). in their study of southern ocean odp sites from maud rise and the georgia basin, mohr & mao (1997) illustrated the taxon manumiella cretacea subsp. gravida, which is similar to some of our specimens. mohr & mao (1997) considered manumiella cretacea subsp. gravida to be an endemic southern ocean taxon and interpreted their assemblages to indicate cooler water conditions. thus, it seems reasonable to assume the same conditions existed in the labrador sea at that time. a similar possible analogy between high southern and northern occurrences has been described by nøhrhansen & dam (1997) from the uppermost maa strichtian strata at nuussuaq, onshore west greenland. here, palynodinium grallator occurs with abundant mio spores and a peak of manumiella sp. that is very similar to manumiella assemblages described from sey mour island, antarctica, by askin (1988) and which she considered were late campanian to late paleocene in age. askin (1988) erected a zone 1, characterised by isabelidinium cretaceum and which she considered cam panian. isabelidinium cretaceum did not occur higher in the seymour island section, being replaced primarily by manumiella species. the occurrence of isabelidinium cretaceum in the labrador sea suggests a mirror-image, high-latitude northern hemisphere re cord. in a recent palynological study of seymour island assemblages, bowman et al. (2012) proposed a formal zonation with two late maastrichtian zones, including two subzones, and one early danian zone. these authors recognised zone 1 of askin (1988), but considered the age to be ?late maastrichtian. bowman et al. (2012) proposed a south polar province for the latest maastrichtian – earliest palaeogene, characterised by species of manu miella, batiacasphaera reticulata and tanyo sphaeridium. again, it seems reasonable for us to consider an equivalent northern hemisphere province during the same time span, although this may have been somewhat warmer than its southern equivalent because of the proximity of the arctic ocean. an al most monospecific assemblage of manumiella see lan dica was reported from seymour island by askin & jacobson (1996) and interpreted to represent marginal marine to shallow-shelf conditions; habib & saeedi (2007) considered manu miella-rich assemblages to re flect cooling during regression at the close of the maa strichtian. determination of palaeogene climates using dino cysts has made considerable advances in recent years (jaramillo & oboh-ikuenobe 1999; sluijs et al. 2005; sluijs et al. 2008; sluijs & brinkhuis 2009). we build on these earlier studies, incorporating taxa that seem to be endemic to midor high latitudes, as noted above and as zonneveld et al. (2013) has shown for modern environments. in the danian, the dinocyst trithyro dinium is common to abundant in the following wells: bjarni o-82, north leif i-05, ogmund e-72 and south labrador n-79 in the hopedale basin (figs 1, 2); gilbert f-53 and skolp e-07 in the saglek basin (figs 1, 2). lentin & williams (1980) included tri thyro dinium in the williams suite for the campanian. nøhr-hansen & dam (1997) and nøhr-hansen et al. (2002) recorded abundant trithyrodinium evittii pulses just above the cretaceous–paleocene boundary at nuus suaq. smit & brinkhuis (1996) and nøhr-hansen & dam (1999) have shown that this species preferred lower latitudes in the late cretaceous but later migrated to higher latitudes, suggesting increasing sea-surface temperatures at high latitudes in the early danian. the abundance of trithyrodinium in higher latitudes was confirmed by nøhr-hansen et al. (2002), who defined an early danian trithyrodinium evittii zone for the succession in the nuussuaq basin, by bowman et al. (2012), who defined an early danian trithyrodinium evittii zone for rocks on seymour island, and by willumsen & vajda (2010) in a study of new zealand dinocyst assemblages. gradstein & srivastava (1980) found that the foraminifera of the labrador margin and nuussuaq reflected atlantic water-mass incursions and that these are supported by mollusc, echinoid and coral faunas on nuussuaq. gradstein & srivastava (1980) considered these observations to indicate tem perate (warmer) climatic conditions in the la brador sea region during the danian; such observations are in accordance with the idea that trithyro di nium testifies to more temperate conditions in the la brador–baffin seaway during the danian. one species that seems to occur consistently in higher latitude samples is palaeocystodinium bulliforme, first described from the paleocene of bylot island by ioan nides (1986). in wells encountering selandian sediments, such as in bjarni o-82 and south labrador n-79 in the hopedale basin and gilbert f-53 and hekja o-71 in the saglek basin (figs 1, 2), palaeo cystodinium bulliforme is common. it is also common in middle to late danian and ?early selandian successions in the nuussuaq basin (nøhr-hansen et al. 2002) and occurs in danian? and selandian sediments on bylot bulletin37.qxp_bulletin 37 26/01/17 13.59 side 45 46 island (g.l. williams, unpublished data). palaeo cystodinium bulliforme is also present in selandian strata in warmer climes farther south, being restricted to the selandian in morocco (h. slimani, personal communication 2015) and having its lo in the uppermost se landian on the scotian margin (fensome et al. 2008). the variation in ranges between regions may be indicative of its origination in more northerly latitudes. the existence of warm climatic conditions in the ypresian and more temperate conditions in the middle eocene (sluijs et al. 2008; zachos et al. 2008; schoon et al. 2013) had an impact on the labrador–baffin sea way, as demonstrated by the presence of nannofossils recovered from the kangâmiut-1 well (sheldon 2003). fur ther evidence for a warm interlude is provided by the influx of the dinocyst apectodinium homomorphum close to the paleocene–eocene boundary in hekja o-71, hellefisk-1, north leif i-05, ogmund e-72 and south labrador n-79. bujak & brinkhuis (1998) con sidered apectodinium homomorphum to be a warmer water spe cies. these findings were confirmed in studies of stable isotope and biogeochemical palaeotemperature indicators (sluijs et al. 2006; zachos et al. 2006; schoon et al. 2013). crouch et al. (2001) showed that the earliest ap pearance of apectodinium-dominated as semblages seems to be synchronous on a global scale. one apectodinium peak occurred during the paleo cene–eocene thermal maximum (petm) at 55 ma, which lasted for about 220 000 years. this peak can be correlated with a negative carbon-isotope excursion (cie), a benthic fora mini fera extinction event and the calcareous nannofossil zo na tion (crouch et al. 2001). ac cording to these authors, the apectodinium influx re flects higher seasurface temperatures and a major in crease in marginal marine surface-water productivity. based on the studies of iakovleva et al. (2001), crouch et al. (2003), sluijs et al. (2008) and others, a marked decline in apectodinium abundances towards the end of the petm may record a corresponding temperature de crease and/or a global regression. it is generally agreed that a maximum flooding surface, defining the culmination of a global transgression, partially coincides with the petm, although the onset of sea-level rise is thought to have predated the petm by a few thousand years. in the labrador–baffin seaway, apecto dinium spp. are common to abundant in the uppermost thanetian in nukik-2 (appendix 3.16), kangâmiut-1 (appendix 3.17), ikermiut-1 (appendix 3.18), hekja o-71 (appendix 3.11), ralegh n-18 (appendix 3.12), gjoa g-37 (appendix 3.13), north leif i-05 (ap pendix 3.1), ogmund e-72 (appendix 3.6) and south labrador n-79 (appendix 3.4). several dinocyst taxa indicate warm-water conditions in the ypresian; examples are species of homo tryblium, which are common in parts of the ypresian section in the following wells: bjarni o-82 (appendix 3.3), north leif i-05 (appendix 3.1) and snorri j-90 (appendix 3.5) in the hopedale basin; gilbert f-53 (appendix 3.9) in the saglek basin; and kangâmiut-1 (appendix 3.17) and nukik-1 (appendix 3.15) in offshore west greenland (figs1, 2). homotryblium is a warm-water genus commonly assumed to favour restricted settings with increased salinity, and thus characterising inshore, lagoonal palaeoenvironments (brink huis 1992, 1994; de verteuil & norris 1996). how ever, in a study of early oligocene dinocyst assemblages from the upper rhine graben of germany, pross & schmiedl (2002) developed a model proposing different settings for what they termed the homotryblium assemblage. they found that the homotryblium assemblage predo minated in nearshore palaeoenvironments where salinity was increased. in these palaeoenvironments, dino cyst species richness was lower. to explain these fluctuations, pross & schmiedl (2002) postulated that dominance of the homotryblium assemblage was related to relatively dry periods with reduced runoff and potentially strong evaporation, leading to high salinity conditions in nearshore palaeoenvironments. ac cor d ing to zonneveld et al. (2013), the related modern species polysphaeridium zoharyi occupies coastal, fully marine, subtropical to tropical regions, which may be characterised by high productivity and high surface-water salinities. specimens of homotryblium are com mon to abundant in the 1890–1930 m interval in gjoa g-37. the presence of the freshwater fern azolla in hekja o71, ralegh n-18 and gjoa g-37 (appendices 3.11 to 3.13) provides further evidence for a warm, humid climate around the ypresian–lutetian boundary. counts for homotryblium and polysphaeridium in the ypresian and their absence from middle eocene and younger rocks indicate that climatic warming was fleeting in the labrador sea. further confirmation is pro vided by the position of the lo of diphyes colli gerum at the lutetian–bartonian boundary in the la brador– baffin seaway. according to brinkhuis & biffi (1993) and bujak & mudge (1994) this species was temperature-sensitive, preferring warmer water palaeo environ ments, explaining its occurrence in the ru pelian in italy. dinocyst species richness takes a dra matic plunge in labrador margin wells in the mid dle eocene. consequently, the numbers of species and specimens decline throughout the remainder of the cenozoic, and the remaining taxa are either ubiquitous or higher-latibulletin37.qxp_bulletin 37 26/01/17 13.59 side 46 47 tude indicators. much of this decline, which was probably related to the general cooling trends during the transition from a greenhouse to an icehouse world (zachos et al. 2001, 2008), mirrors the global decline in dinocyst species richness during the cenozoic, as shown by macrae et al. (1996). dinocyst assemblages from the eocene–oligocene strata of odp site 647a in the southern labrador sea (figs 1, 2) provide some clues to palaeogene oceanic conditions. from a study of odp assemblages, head & norris (1989) concluded that evidence existed for a proto-gulf stream in the middle eocene. the odp assemblages included taxa known only from the southwestern atlantic, eastern united states, norwegian sea, belgium and australia. such a distribution could be best explained by the pattern of oceanic currents. a cooling trend in the late eocene – oligocene is indicated by an influx of colder-water taxa, including gelatia inflata, svalbardella and protoperidiniaceans (head & norris 1989). this cooling trend in the late eocene – early oligocene is also a characteristic of high latitudes in the southern hemisphere (guerstein et al. 2008, 2010; houben et al. 2013) and probably reflects cooling conditions in the oligocene, resulting from a reduction in atmospheric pco2, which allowed develop ment of a permanent ice sheet in antarctica (hren et al. 2013). the relative decline in dinocyst species richness in labrador margin wells is accompanied, especially in the oligocene and neogene, by a drop in angiosperm po l len species richness (williams 1986). however, conifer pollen in the neogene, primarily pinuspollenites, are common. general trends suggest rapid and widespread cooling for the labrador sea, triggered in part by fluctuations in oceanic-circulation patterns. at odp site 646, also in the labrador sea (figs 1, 2), the late miocene – early pliocene dinocyst associations reflect predominantly temperate to cool surface waters, with impagidinium pallidum and diverse protoperidiniaceans (head et al. 1989a). zonneveld et al. (2013) noted that impagidi nium pallidum is primarily restricted to higher latitudes with high concentrations in arctic and antarctic re gions, whereas of the two species of protoperidinium, p. americanum is a coastal sub– polar to tropical form whereas p. monospinum is characteristically found in “full marine, tropical to equa torial upwelling areas of nw africa” (zonneveld et al. 2013, p. 131). these observations could indicate mixing of warm-temperate and arctic water in the labrador sea, with the west greenland current al ready in place – a situation not dissimilar to present-day conditions. according to de vernal & mudie (1989a), the early miocene assemblages at odp site 645 in baffin bay (figs 1, 2) seem to indicate cool-temperate surface waters. de vernal & mudie (1989a, b) found that similar conditions persisted into the pliocene–pleistocene at both site 646 and site 647 (figs 1, 2). there was a marked drop in species richness in the middle and late miocene, however, accompanied by an increase in ter rigenous debris. head et al. (1989b) related these changes to the onset of colderwater conditions and the establishment of circulation pat terns in baffin bay similar to those of today. however, dapsilidinium pa stielsii, one of the species recorded by these authors, was considered to be a warmer water species by head & westphal (1999), who noted that its latitudinal occurrences contracted in the late miocene and plio cene; they attributed this to the cooling of the north at lantic in the late miocene and the evolution of the cold la brador current. the persistence of dapsilidinium pa stielsii into the middle miocene is probably a reflection of the midmiocene climatic optimum (zachos et al. 2008). piasecki (2003) concluded that the dinocyst assemblages in qulleq-1 on the west greenland margin (figs 1, 2; appendix 3.14) were comparable to north atlantic assemblages, but had higher latitude elements. the ser ravallian samples reflected warmer water, signifying the mid-miocene climatic optimum. in contrast, the plio cene was marked by depleted dinocyst counts re flecting the progressively deteriorating climate. schreck et al. (2012, 2013), in their study of the dino cyst assemblages from an almost continuous mid dle miocene through pliocene section in odp site 907a in the iceland sea, recorded high marine productivity in the middle miocene, with a diminution toward the impoverished assemblages of the late pliocene. we con clude that the proto-east greenland current was well developed by about 9 ma. the establishment at around 4.5 ma of the modern east greenland current, which would presumably have accelerated the influence and impact of the labrador current, may explain the low abundance and species richness of the dinocyst assamblages. de vernal & mudie (1989a) determined that the pliocene–pleistocene dinocysts at site 645 (figs 1, 2) included boreal and cool-temperate taxa. these data suggest that the major cooling occurred relatively late in the cenozoic history of baffin bay. the waning influence of the gulf stream and the development of the labrador current were important events in the evolution of the labrador sea. based on foraminiferal data from a study of dsdp sites 111 and 112, poore & berggren (1974) considered the cut-off of the gulf stream to have occurred in the middle–late pliocene. bulletin37.qxp_bulletin 37 26/01/17 13.59 side 47 48 conclusions a new biostratigraphic framework has been developed for the aptian (lower cretaceous) to pliocene–pleisto cene of the labrador–baffin seaway. this framework is based on 187 bioevents for taxa of dinocysts, mio spores, fungal spores and azolla; these define 106 bio event horizons, mostly last occurrences, but including some local and regional peakor common-occurrence events. also incorporated are additional data from pre vious studies by piasecki (2003) and pedersen & nøhrhansen (2014). most events are concentrated in the cam panian to rupelian interval. integration of the palynological data from wells on both margins of the seaway provides the first broad biostratigraphic correlation of mesozoic–cenozoic strata of the region. detailed biostratigraphic evidence has confirmed the following hiatuses: pre-aptian in the hopedale basin; pre-albian in the saglek basin; albian–turonian in some wells of the hopedale basin; turonian–santo nian/campanian in some areas; pre-campanian and late campanian–thanetian on the greenland margin; late maastrichtian and danian in some wells of the hopedale basin and in the saglek basin; selandian in part in the hopedale basin, in total in the saglek basin and in two wells on the west greenland margin; late ypresian and/or lutetian on both sides of the seaway; oligocene to middle miocene of considerable variability on both margins, with all of the oligocene and the lower miocene missing in all the west greenland mar gin wells; middle to late miocene on the western side (figs 7–9). on the canadian margin, these hiatuses can be matched in part with the five recognised regional unconformities of mcwhae (1981) as follows: the preaptian–albian hiatus of this study is presumed to cor relate with the labrador unconformity; the aptian/ albian–turonian hiatus represents the avalon un conformity; the bylot unconformity is represented by the maastrichtian–selandian hiatus and the ‘base-ter tiary unconformity’ in the jeanne d’arc basin (sinclair 1988); the oligocene to middle miocene hiatus of this study may be equivalent to the baffin bay uncon formity but we regard it as resulting from the late oligocene drop in sea level, which is not mentioned by mcwhae et al. (1980); and finally the middle to late miocene hiatus described here matches the proposed age of the beaufort unconformity. mcwhae (1981) only recognised these five unconformities and restricted them to specific ages. this does not fully match our findings since we recognise seven significant hiatuses which furthermore are of greater duration than indicated by mcwhae (1981). the intra-selandian and the lutetian hiatuses documented here are not readily correlated with the regional unconformities of mcwhae (1981), but may match additional seismic unconformities reported by dickie et al. (2011). on the west greenland margin, the hiatuses are fewer and only partly conform to those on the canadian margin. the absence of a pre-santonian section in part reflects the depth of the wells, since seismic data indicate that older strata are present. an apparent unconformity separating the appat seismic sequence from the overlying kangeq seismic sequence, probably equates with the avalon unconformity. the late campanian to selan dian hiatus in two wells and the absence of the danian and most or all of the selandian in the other four wells represent the bylot unconformity. thus it appears that the bylot unconformity is equivalent to the ‘base-ter tiary unconformity’ recognised by dalhoff et al. (2003). alternatively, following dickie et al. (2011), the bylot unconformity could be restricted to the selandian. the intra-selandian hiatus recognised here, however, demonstrates the difficulty of identifying the bylot un con formity sensu stricto. on the west greenland margin, dal hoff et al. (2003) also mapped a mid eocene un con formity, which we consider equivalent to our lu tetian hiatus. in three of the wells, the overlying se di ments are middle miocene in age. in two other wells the bartonian is overlain by middle miocene sediments and in one the priabonian is overlain by middle miocene sediments. these hiatuses could be regarded as correlating with the baffin bay unconformity, but the time spans are significantly greater than that recorded by mcwhae (1981) on the canadian margin, and they are likely to represent composite unconformities. palynomorph assemblages show that most of the ap tian and albian sediments on the present-day labrador margin were deposited in marginal marine to lagoonal palaeoenvironments, punctuated by a shallow marine episode in the aptian. a marine transgression starting in the cenomanian–turonian led to the most offshore, presumably deepest water palaeoenvironments in the campanian – late maastrichian, although this conclusion seems to conflict with the hiatus between the cre taceous and cenozoic sediments in many wells, espe cially on the west greenland margin. however, the cretaceous–pa bulletin37.qxp_bulletin 37 26/01/17 13.59 side 48 49 laeo gene boundary and upper maastrichtian and da n ian–selandian strata occur in the nuussuaq basin (nøhrhansen & dam 1997; dam et al. 2009). outer neritic to open-ocean conditions persisted throughout the paleocene and ypresian and into the lutetian, an interval corresponding to a time of sea-floor spreading in the labrador–baffin seaway. the onset of shallowing, probably in the late lutetian, continued through the pria bonian and into the rupelian, when marginal marine to inner neritic palaeoenvironments predominated. throug h out the rest of the oligocene and in the neogene, inner neritic palaeoenvironments alternated with marginal marine conditions. dinocysts indicate that climatic conditions in the labrador–baffin seaway region, which had been relatively temperate in the cretaceous, varied dramatically in the cenozoic. the palaeogene was a time of increasingly warmer climate, a thermal maximum being reached around the paleocene–eocene boundary, reflecting the global thermal event at this time. warm to hot conditions prevailed throughout the ypresian, but began to cool in the lutetian and cooling accelerated in the pria bonian and rupelian, a trend observed globally (zachos et al. 2008). temperatures generally declined throughout the neogene, reaching a low in the pleistocene. acknowledgements during the gestation of this paper of over a decade, we have had the unwavering support of our two institutions, the geological survey of denmark and greenland (geus) and the geological survey of canada (atlantic), (gsca), part of the earth sciences sector (ess) of natural resources canada. we are also grateful for fruitful discussions with and feedback from kate dickie of gsca, lotte m. larsen, gunver k. pedersen and martin sønderholm of geus, and stefan piasecki of the geological museum, natural history mu seum of denmark, university of copenhagen. we are grateful to lynn dafoe and christop her harrison for incisive reviews of an earlier draft that led to important improvements of the manuscript. we thank the two bulletin referees, raquel guerstein and gregers dam for their very valuable com ments and suggestions. we also thank bernie cri lley and bill macmillan (both gsc) and annette ryge and dorthe salomonsen (both geus) for processing samples; nelly koziel (gsc) for technical support; and lynn dafoe (gsc) for collecting and providing con ventional core samples. this paper is ess contribution number 20150377. bulletin37.qxp_bulletin 37 26/01/17 13.59 side 49 50 references ainsworth, n.r., riley, l.a., bailey, h.w. & gueinn, k.j. 2014: cretaceous – tertiary stratigraphy of the labrador shelf, 56 pp. riley geoscience ltd. andrews, j.t., wray, j.l., guennel, g.k. & ives, j.d. 1972: an early tertiary outcrop in north-central baffin island, northwest territories, canada: environment and significance. canadian journal of earth sciences 9 (3), 233–238. askin, r.a. 1988: campanian to paleocene palynological succession of seymour and adjacent islands, northeastern antarctic peninsula. in: feldmann, r.m. & woodburne, m.o. 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(1990) reference in zonneveld et al. (2013) reference in fensome et al. (2016) reference in http://botany.si.edu/ing/ reference in http://fossilworks.org all others can be found in http://dinoflaj.smu.ca/wiki/ bulletin37.qxp_bulletin 37 26/01/17 13.59 side 60 61 atopodinium haromense thomas & cox 1988 axiodinium augustum (harland 1979c) williams et al. 2015 balmula tripenta bint 1986 (now nyktericysta tripenta) barssidinium evangelineae lentin et al. 1994 barssidinium graminosum lentin et al. 1994 batiacasphaera reticulata (davey 1969b) davey 1979d batioladinium jaegeri (alberti 1961) brideaux 1975 callaiosphaeridium asymmetricum (deflandre & courteville 1939) davey & williams 1966b cannosphaeropsis wetzel 1933b cannosphaeropsis passio de verteuil & norris 1996a cauveridinium membraniphorum (cookson & eisenack 1962b) masure in fauconnier & masure 2004 cerbia tabulata (davey & verdier 1974) below 1981a cerebrocysta bartonensis bujak in bujak et al. 1980 cerebrocysta magna bujak 1994 cerebrocysta poulsenii de verteuil & norris 1996a cerodinium vozzhennikova 1963 cerodinium diebelii (alberti 1959b) lentin & williams 1987 cerodinium glabrum (gocht 1969) fensome et al. 2009 cerodinium kangiliense nøhr-hansen & heilmann-clausen 2001 cerodinium pannuceum (stanley 1965) lentin & williams 1987 cerodinium speciosum (alberti 1959b) lentin & williams 1987 cerodinium speciosum subsp. glabrum (gocht 1969) lentin & williams 1987 (now cerodinium glabrum) charlesdowniea columna (michoux 1988) lentin & vozzhennikova 1990 (now piladinium columnum) chatangiella (vozzhennikova 1967) fensome et al. 2016 * chatangiella decorosa (mcintyre 1975) lentin & williams 1976 chatangiella ditissima (mcintyre 1975) lentin & williams 1976 chatangiella madura lentin & williams 1976 chatangiella mcintyrei nøhr-hansen 1996 chiropteridium galea (maier 1959) sarjeant 1983 chiropteridium gilbertii fensome et al. 2016 * chlamydophorella nyei cookson & eisenack 1958 chytroeisphaeridia hadra fensome et al. 2016 * circulodinium distinctum (deflandre & cookson 1955) jansonius 1986 cleistosphaeridium davey et al. 1966 cleistosphaeridium ancyreum (cookson & eisenack 1965a) eaton et al. 2001 cleistosphaeridium diversispinosum davey et al. 1966 cleistosphaeridium palmatum fensome et al. 2016 * cleistosphaeridium placacanthum (deflandre & cookson 1955) eaton et al. 2001 cleistosphaeridium polypes (cookson & eisenack 1962b) davey 1969a (now kiokansium unituberculatum) cleistosphaeridium polypetellum (islam 1983c) stover &williams 1995 cordosphaeridium eisenack 1963b cordosphaeridium cantharellus (brosius 1963) gocht 1969 cordosphaeridium delimurum fensome et al. 2009 cordosphaeridium fibrospinosum davey & williams 1966b cordosphaeridium funiculatum morgenroth 1966a cordosphaeridium gracile (eisenack 1954b) davey & williams 1966b cordosphaeridium minimum (morgenroth 1966a) benedek 1972 (now minisphaeridium latirictum) cribroperidinium neale & sarjeant 1962 cribroperidinium giuseppei (morgenroth 1966a) helenes 1984 cyclonepheliium attadalicum cookson & eisenack 1962b dapsilidinium pastielsii (davey & williams 1966b) bujak et al. 1980 bulletin37.qxp_bulletin 37 26/01/17 13.59 side 61 62 dapsilidinium pseudocolligerum (stover 1977) bujak et al. 1980 dapsilidinium pseudoinsertum fensome et al. 2016 * deflandrea eisenack 1938b deflandrea galeata (lejeune-carpentier 1942) lentin & williams 1973 deflandrea majae (schiøler 1993) fensome et al. 2016 * deflandrea oebisfeldensis alberti 1959b deflandrea phosphoritica eisenack 1938b dinogymnium evitt et al. 1967 dinogymnium longicorne (vozzhennikova 1967) harland 1973 dinogymnium sibiricum (vozzhennikova 1967) lentin & williams 1973 diphyes brevispinum bujak 1994 diphyes colligerum (deflandre & cookson 1955) cookson 1965a diphyes ficusoides islam 1983b disphaerogena carposphaeropsis wetzel 1933b eatonicysta furensis (heilmann-clausen in heilmann-clausen & costa 1989) stover & williams 1995 eatonicysta ursulae (morgenroth 1966a) stover & evitt 1978 edwardsiella sexispinosa versteegh & zevenboom in versteegh 1995 endoceratium dettmanniae (cookson & hughes 1964) stover & evitt 1978 enneadocysta magna fensome et al. 2007 eocladopyxis morgenroth 1966a eocladopyxis peniculata morgenroth 1966a evittosphaerula? foraminosa fensome et al. 2016 * fibrocysta bipolaris (cookson & eisenack 1965b) stover & evitt 1978 gelatia inflata bujak 1984 geonettia de verteuil & norris 1996b gillinia hymenophora cookson & eisenack 1960a ginginodinium? flexidentatum fensome et al. 2016 * glaphyrocysta stover & evitt 1978 glaphyrocysta divaricata (williams & downie 1966c) stover & evitt 1978 glaphyrocysta exuberans (deflandre & cookson 1955 ex eaton 1976) stover & evitt 1978 glaphyrocysta retiintexta (cookson 1965a) stover & evitt 1978 glaphyrocysta spineta (eaton 1976) stover & evitt 1978 glaphyrocysta texta (bujak 1976) stover & evitt 1978 glaphyrocysta vicina (eaton 1976) stover & evitt 1978 gonyaulacysta fastigiata duxbury 1977 gonyaulacysta pectinigera (gocht 1970b) fensome 1979 habibacysta tectata head et al. 1989b hapsocysta? benteae nøhr-hansen 1993 heteraulacacysta drugg & loeblich jr. 1967 heteraulacacysta porosa bujak in bujak et al. 1980 heterosphaeridium cookson & eisenack 1968 heterosphaeridium bellii radmacher et al. 2014 heterosphaeridium difficile (manum & cookson 1964) ioannides 1986 heterosphaeridium heteracanthum (deflandre & cookson 1955) eisenack & kjellström 1972 homotryblium davey & williams 1966b homotryblium abbreviatum eaton 1976 homotryblium tenuispinosum davey & williams 1966b hurlandsia rugara (piasecki 1984) lister & batten 1988a hystrichokolpoma klumpp 1953 hystrichokolpoma bulbosum (ehrenberg 1838) morgenroth 1968 hystrichokolpoma rigaudiae deflandre & cookson 1955 bulletin37.qxp_bulletin 37 26/01/17 13.59 side 62 63 hystrichosphaeridium deflandre 1937b hystrichosphaeridium quadratum fensome et al. 2016 * hystrichosphaeridium tubiferum (ehrenberg 1838) deflandre 1937b hystrichosphaeropsis perforata schiøler 1993 hystrichosphaeropsis quasicribrata (wetzel 1961) gocht 1976 impagidinium stover & evitt 1978 impagidinium dispertitum (cookson & eisenack 1965a) stover & evitt 1978 impagidinium victorianum (cookson & eisenack 1965a) stover & evitt 1978 impagidinium pallidum bujak 1984 impletosphaeridium apodastum fensome et al. 2016 * invertocysta lacrymosa edwards 1984 isabelidinium lentin & williams 1977a isabelidinium cooksoniae (alberti 1959b) lentin & williams 1977a isabelidinium cretaceum (cookson 1956) lentin & williams 1977a isabelidinium magnum (davey 1970) stover & evitt 1978 isabelidinium microarmum (mcintyre 1975) lentin & williams 1977a kiokansium unituberculatum (tasch in tasch et al. 1964) stover & evitt 1978 kiokansium williamsii singh 1983 kleithriasphaeridium mantellii (davey & williams 1966b) fensome et al. 2016 * labyrinthodinium truncatum piasecki 1980 laciniadinium mcintyre 1975 laciniadinium arcticum (manum & cookson 1964) lentin & williams 1980 lejeunecysta artzner and dörhöfer 1978 lentinia serrata bujak in bujak et al. 1980 licracysta corymbus fensome et al. 2007 licracysta? semicirculata (morgenroth 1966b) fensome et al. 2007 lingulodinium wall 1967 lithodinia eisenack 1935 manumiella bujak & davies 1983 manumiella cretacea subsp. gravida (mao shaozhi & mohr 1992) lentin & williams 1993 manumiella seelandica (lange 1969) bujak & davies 1983 mendicodinium morgenroth 1970 microdinium ornatum cookson & eisenack 1960a minisphaeridium latirictum (davey & williams 1966b) fensome et al. 2009 nelchinopsis kostromiensis (vozzhennikova 1967) wiggins 1972 nelsoniella cookson & eisenack 1960a nematosphaeropsis deflandre & cookson 1955 nyktericysta bint 1986 nyktericysta arachnion bint 1986 nyktericysta davisii bint 1986 nyktericysta dictyophora he chengquan et al. 1992 nyktericysta tripenta (bint 1986) fensome et al. 2009 odontochitina ancala bint 1986 odontochitina costata alberti 1961 odontochitina operculata (wetzel 1933a) deflandre & cookson 1955 odontochitina porifera cookson 1956 oligosphaeridium albertense (pocock 1962) davey & williams 1969 oligosphaeridium totum brideaux 1971 operculodinium wall 1967 operculodinium centrocarpum (deflandre & cookson 1955) wall 1967 operculodinium giganteum wall 1967 bulletin37.qxp_bulletin 37 26/01/17 13.59 side 63 64 operculodinium janduchenei head et al. 1989b operculodinium piaseckii strauss & lund 1992 palaeocystodinium australinum (cookson 1965b) lentin & williams 1976 palaeocystodinium bulliforme ioannides 1986 palaeocystodinium golzowense alberti 1961 palaeohystrichophora infusorioides deflandre 1935 palaeoperidinium deflandre 1934 ex sarjeant 1967b palaeoperidinium pyrophorum (ehrenberg 1838 ex wetzel 1933a) sarjeant 1967b palynodinium grallator gocht 1970a perisseiasphaeridium pannosum davey & williams 1966b petalodinium condylos (williams & downie 1966b) williams et al. 2015 phelodinium stover & evitt 1978 phelodinium kozlowskii (górka 1963) lindgren 1984 phthanoperidinium drugg & loeblich jr. 1967 phthanoperidinium alectrolophum eaton 1976 phthanoperidinium coreoides (benedek 1972) lentin & williams 1976 phthanoperidinium geminatum bujak in bujak et al. 1980 phthanoperidinium levimurum bujak in bujak et al. 1980 phthanoperidinium multispinum bujak in bujak et al. 1980 phthanoperidinium regale bujak 1994 phthanoperidinium stockmansii (de coninck 1975) lentin & williams 1977b piladinium columna (michoux 1988) williams et al. 2015 polysphaeridium davey & williams 1966b polysphaeridium zoharyi (rossignol 1962) bujak et al. 1980 protoperidinium bergh 1881a protoperidinium americanum (paulsen 1907) zonneveld & dale 1994 ^^ protoperidinium monospinum (gran & braarud 1930) balech 1974 ^^ pseudoceratium gocht 1957 pseudoceratium interiorense bint 1986 pterodinium eisenack 1958a pyrophacus stein 1883 raphidodinium fucatum deflandre 1936b reticulatosphaera actinocoronata (benedek 1972) bujak & matsuoka 1986 rhombodinium draco gocht 1955 rhombodinium porosum bujak 1979 scalenodinium scalenum fensome et al. 2016 * schematophora speciosa deflandre & cookson 1955 scriniodinium obscurum manum & cookson 1964 (now spongodinium obscurum) selenopemphix brevispinosa head et al. 1989c selenopemphix nephroides benedek 1972 senegalinium jain & millepied 1973 senegalinium iterlaaense nøhr-hansen & heilmann-clausen 2001 senoniasphaera inornata (drugg 1970b) stover & evitt 1978 senoniasphaera microreticulata brideaux & mcintyre 1975 senoniasphaera protrusa clarke & verdier 1967 senoniasphaera rotundata clarke & verdier 1967 simplicidinium insolitum (eaton 1976) fensome et al. 2016 * sirmiodinium grossii alberti 1961 sophismatia tenuivirgula (williams & downie 1966b) williams et al. 2015 spinidinium cookson & eisenack 1962b spinidinium densispinatum stanley 1965 bulletin37.qxp_bulletin 37 26/01/17 13.59 side 64 65 spinidinium echinoideum (cookson & eisenack 1960a) lentin & williams 1976 spiniferites mantell 1850 spiniferites ovatus matsuoka 1983b spiniferites pseudofurcatus (klumpp 1953) sarjeant 1970 spiniferites scabrosus (clarke & verdier 1967) lentin & williams 1975 spiniferites solidago de verteuil & norris 1996a (=achomosphaera grallaeformis (brosius 1963) davey & williams 1969) spongodinium delitiense (ehrenberg 1838) deflandre 1936b spongodinium grossum (manum & cookson 1964) fensome et al. 2016 * spongodinium obscurum (manum & cookson 1964) fensome et al. 2016 * stichodinium lineidentatum (deflandre & cookson 1955) williams et al. 2015 stiphrosphaeridium dictyophorum (cookson & eisenack 1958) lentin & williams 1985 subtilisphaera jain & millepied 1973 subtilisphaera perlucida (alberti 1959b) jain & millepied 1973 surculosphaeridium longifurcatum (firtion 1952) davey et al. 1966 svalbardella manum 1960 systematophora ancyrea cookson & eisenack 1965a (now cleistosphaeridium ancyreum) talladinium pellis fensome et al. 2016 * tanyosphaeridium davey & williams 1966b tanyosphaeridium xanthiopyxides (wetzel 1933b ex deflandre 1937b) stover & evitt 1978 taurodinium granulatum fensome et al. 2016 * tenua hystrix eisenack 1958a thalassiphora fenestrata liengjarern et al. 1980 trichodinium castanea deflandre 1935 ex clarke & verdier 1967 trinovantedinium reid 1977 trithyrodinium drugg 1967 trithyrodinium? conservatum fensome et al. 2016 * trithyrodinium evittii drugg 1967 trithyrodinium quinqueangulare marheinecke 1992 trithyrodinium suspectum (manum & cookson 1964) davey 1969b tuberculodinium wall 1967 tuberculodinium vancampoae (rossignol 1962) wall 1967 vesperopsis bint 1986 vesperopsis longicornis (batten & lister 1988) harding 1990b vesperopsis mayi bint 1986 vesperopsis nebulosa bint 1986 wallodinium luna (cookson eisenack 1960a) lentin & williams 1973 wetzeliella eisenack 1938b emend. williams, damassa, fensome & guerstein in fensome et al. 2009 wetzeliella homomorpha deflandre & cookson 1955 (now apectodinium homomorphum) wetzeliella parva alberti 1961 (now apectodinium parvum) xenascus ceratioides (deflandre 1937b) lentin & williams 1973 xenascus sarjeantii (corradini 1973) stover & evitt 1978 acritarchs and algae cymatiosphaera invaginata head et al. 1989a ** fromea nicosia jansonius 1989 * fromea quadrangularis fensome et al. 2016 * fromea tornatilis (drugg 1978) lentin & williams 1981 micrhystridium deflandre 1937b microsphaeridium ancistroides benedek 1972 paralecaniella indentata (deflandre & cookson 1955) cookson & eisenack 1970b bulletin37.qxp_bulletin 37 26/01/17 13.59 side 65 66 pediastrum meyen 1829 * tetraporina naumova 1939 plant microfossils afropollis doyle et al. 1982 * alisporites grandis (cookson 1953) dettmann 1963 # appendicisporites jansonii pocock 1962 # aquilapollenites rouse 1957 * artemisia linneaus 1753 § azolla lamarck in lamarck et al. 1783 balmeisporites holodictyus cookson & dettmann 1958 † bombacacidites couper 1960 † callialasporites dampieri (balme 1957) dev 1961 * callialasporites obrutus norris 1969 * caryapollenites raatz 1938 ex potonié 1960 * cerebropollenites mesozoicus (couper 1958) nilsson 1958 * cicatricosisporites australiensis (cookson 1953) potonié 1956 # cicatricosisporites minutaestriatus (bolkhovitina 1961) pocock 1964 * cicatricosisporites ornatus srivastava 1972 * cicatricosisporites reticicatricosus döring 1965a > cicatricososporites eocenicus (selling 1944) jansonius & hills 1976 * compositoipollenites potonié 1951 ex potonié 1960 * contignisporites glebulentus dettmann 1963 # corsinipollenites oculusnoctis (thiergart 1940) nakoman 1965 * extratriporopollenites pflug in thomson & pflug 1952 ex pflug in thomson & pflug 1953 * graminidites cookson 1947 ex potonié 1960 * klukisporites areolatus singh 1971 # osmundacidites wellmannii couper 1953 * parvisaccites amplus brenner 1963 * parvisaccites radiatus couper 1958 * periporopollenites pflug & thomson in thomson & pflug 1953 * pilosisporites verus delcourt & sprumont 1955 *** pinus linnaeus 1753 § pinuspollenites raatz 1938 ex potonié 1958 † pistillipollenites macgregorii rouse 1962 * plicatella bifurcata (singh 1964) dörhöfer 1977 > quercoidites potonié et al. 1950 ex potonié 1960 * rugubivesiculites pierce 1961 ^ rugubivesiculites convolutus pierce 1961 † rugubivesiculites multiplex pierce 1961 ^ rugubivesiculites multisaccus singh 1983 †† rugubivesiculites reductus pierce 1961 ^ rugubivesiculites rugosus pierce 1961 ^ taraxacum wiggers 1780 § tiliaepollenites crassipites (wodehouse 1933) fensome et al. 2016 * tsugaepollenites igniculus potonié 1931 * (now zonalapollenites igniculus) wodehouseia spinata stanley 1961 * zlivisporis pacltová 1961 * zonalapollenites igniculus (potonié 1931) thomson & pflug 1953 * bulletin37.qxp_bulletin 37 26/01/17 13.59 side 66 67 palaeoenvironmental analysis of six wells on the canadian margin (for location, see fig. 1), based on palaeoecolo gical interpretation of palynological data (by graham l. williams). legend for appendices 2.1–2.6 appendix 2 cuttings sample cuttings sample (secondary suite) side-wall core (swc) bulletin37.qxp_bulletin 37 26/01/17 13.59 side 67 68 pliocene – pleistocene serravallian early miocene chattian rupelian priabonian bartonian 405–415 430–440 465–475 490–500 525–535 555–565 585–595 615–625 645–655 675–685 700–710 735–745 765–775 790–800 825–835 855–865 885–895 950–960 970–980 1035–1045 1065–1075 1095–1105 1125–1135 1155–1165 1215–1225 1275–1285 1305–1315 1335–1345 1365–1375 1395–1405 1425–1435 1455–1465 1485–1495 depositional environment depositional environment n o n -m ar in e c o as ta l – m ar gi n al m ar in e in n er n er it ic o u te r n er it ic o p en o ce an sample depth (m)age period/ epoch age period/ epoch sample depth (m) 400 (m) 500 600 700 800 900 1000 1100 1200 1300 1400 1500 lutetian ypresian bartonian basal ypresian danian selandian maastrichtian e. campanian turonian– coniacian early albian cenomanian? aptian 1525–1535 1545–1555 1575–1585 1605–1615 1635–1645 1660–1670 1725–1735 1695–1705 1755–1765 1785–1795 1815–1825 1875–1885 1905–1915 1935–1945 1965–1975 1995–2005 2025–2035 2055–2065 2085–2095 2115–2125 2145–2155 2175–2185 2205–2215 2235–2245 2265–2275 2295–2305 2350–2360 2330–2340 2380–2390 2415–2425 2445–2455 2475–2485 2535–2545 2565–2575 2595–2605 2625–2635 2645–2655 1500 1600 1700 1800 1900 2000 2100 2200 2300 2400 2500 2600 n o n -m ar in e c o as ta l – m ar gi n al m ar in e in n er n er it ic o u te r n er it ic o p en o ce an m o c e n e o l ig o c e n e e o c e n e e l e l m m e o c e n e p a l e o c e n e c r e t a c e o u s m e l e l e appendix 2.1. stratigraphy and palaeoenvironmental interpretation of the bjarni o-82 well. kelly bushing height: 12.0 m above sea level. water depth: 144 m. total depth: 2650.0 m. bulletin37.qxp_bulletin 37 26/01/17 13.59 side 68 69 pliocene– pleistocene middle miocene m o c e n e o l ig o c e n e e o c e n e rupelian ypresian early miocene priabonian bartonian early lutetian 720–730 750–760 780–790 810–820 840–850 870–880 900–910 930–940 960–970 990–1000 1020–1030 1050–1060 1080–1090 1110–1120 1140–1150 1170–1180 1200–1210 1230–1240 1260–1270 1290–1300 1320–1330 1350–1360 1380–1390 1410–1420 1440–1450 1470–1480 1500–1510 1530–1540 1560–1570 1590–1600 1620–1630 1650–1660 1680–1690 1710–1720 1740–1750 1770–1780 1800–1810 1920–1930 1950–1960 1980–1990 2010–2020 2040–2050 2070–2080 2100–2110 2130–2140 1830–1840 1890–1900 800 (m) 900 1000 1100 1200 1300 1400 1500 1600 1700 1800 1900 2000 2100 thanetian danian aptian precambrian early ypresian selandian early maastrichtian campanian coniacian 2130–2140 2160–2170 2190–2200 2220–2230 2250–2260 2280–2290 2310–2320 2340–2350 2370–2380 2400–2410 2430–2440 2460–2470 2490–2500 2550–2560 2580–2590 2610–2620 2640–2650 2700–2710 2730–2740 2760–2770 2790–2800 2820–2830 2850–2860 2520–2530 2880–2890 2910–2920 2940–2950 2970–2980 3000–3010 3030–3040 3060–3070 3090–3100 3120–3130 3150–3160 3180–3190 3210–3220 3240–3250 3270–3280 3300–3310 3330–3340 3360–3370 3390–3400 3420–3430 3450–3460 3480–3490 3510–3520 3540–3550 3545–3555 3571.5 3565–3571.5 2200 (m) 2300 2400 2500 2600 2700 2800 2900 3000 3100 3200 3300 3400 3500 depositional environment n o n -m ar in e c o as ta l – m ar gi n al m ar in e m ar gi n al m ar in e in n er n er it ic o u te r n er it ic o p en o ce an depositional environment n o n -m ar in e c o as ta l – in n er n er it ic o u te r n er it ic o p en o ce an age period/ epoch age period/ epoch sample depth (m) sample depth (m) ypresian m e e e l m e e e l l e o c e n e p a l e o c e n e c r e t a c e o u s appendix 2.2. stratigraphy and palaeoenvironmental interpretation of the south labrador n-79 well. kelly bushing height: 11.3 m above sea level. water depth: 449.9 m. total depth: 3571.5 m. bulletin37.qxp_bulletin 37 26/01/17 13.59 side 69 70 pliocene– pleistocene pliocene late miocene 384.05–393.19 414.53–423.67 445.01–454.15 475.49–484.63 505.97–515.11 536.45–545.59 566.93–576.07 597.41–606.55 627.89–637.03 658.37–667.51 688.85–697.99 719.33–728.47 749.81–758.95 780.29–789.43 810.77–819.91 841.25–850.39 932.69–941.83 950.98–960.12 978.41–987.55 1005.84–1014.98 1033.27–1042.42 1124.71–1133.86 1152.14–1161.29 1261.87–1271.02 1295.4 1316.74–1325.88 1344.17–1353.31 1371.6 1399.03–1408.18 1426.46–1435.61 1453.90–1463.04 1481.33–1490.47 1508.76–1517.91 1536.19–1545.34 1563.62–1572.77 1591.06–1597.16 1618.49–1627.63 1645.92 1673.35–1682.50 1700.79–1709.93 1728.22–1737.37 1755.65–1764.80 1776.98 1783.09–1792.23 400 (m) 500 600 700 800 900 1000 1100 1200 1300 1400 1500 1600 1700 1800 bartonian lutetian ypresian basal ypresian thanetian selandian danian aptian precambrian 1810.52–1819.66 1837.95–1847.09 1865.38–1874.53 1892.81–1901.96 1920.25–1929.39 1947.68–1956.82 1975.11–1984.25 2005.59–2014.73 2029.97–2039.12 2072.64 2084.83–2093.98 2112.27–2121.41 2139.70–2148.85 2161.03 2194.57–2203.71 2240.29 2249.43–2258.57 2293.63 2304.29–2313.44 2340.87 2359.16–2368.30 2386.59–2395.74 2420.12 2441.46–2450.60 2468.89–2478.03 2497.23–2497.54 2499.37–2499.67 2520.70 2551.18–2560.33 2578.62–2587.76 2609.10–2618.24 2636.53–2645.67 2663.96–2673.10 2700.53 2724.91 2763.32 2779.78–2788.93 2820.02 2840.74–2849.89 2871.22–2880.37 2889.51 2915.42 2932.18–2941.33 2962.66–2971.81 2993.15–3002.29 3025.45 3037.34 3054.11–3063.25 3127.56 3145.55–3154.69 3176.03–3185.17 3200.41–3209.55 1800 1900 2000 2100 2200 2300 2400 2500 2600 2700 2800 2900 3000 3100 3200 depositional environment n o n -m ar in e c o as ta l – m ar gi n al m ar in e in n er n er it ic o u te r n er it ic o p en o ce an depositional environment n o n -m ar in e c o as ta l – m ar gi n al m ar in e in n er n er it ic o u te r n er it ic o p en o ce an age period/ epoch age period/ epoch sample depth (m) sample depth (m) early–middle miocene early oligocene early priabonian early priabonian p l io c e n e – p l e is t o c e n e m o c e n e o l ig o c e n e e o c e n e p l io c e n e l l e o c e n e p a l e o c e n e c r e t a c e o u s l m e e e l e e-m appendix 2.3. stratigraphy and palaeoenvironmental interpretation of the snorri j-90 well. kelly bushing height: 11.3 m above sea level. water depth: 140.8 m. total depth: 3209.8 m. bulletin37.qxp_bulletin 37 26/01/17 14.00 side 70 71 late miocene early–middle miocene chattian rupelian priabonian bartonian lutetian ypresian 525–535 555–565 585–595 615–625 645–655 675–685 705–715 735–745 765–775 795–805 825–835 855–865 885–895 915–925 945–955 975–985 1005–1015 1035–1045 1065–1075 1095–1105 1125–1135 1155–1165 1185–1195 1215–1225 1245–1255 1275–1285 1305–1315 1335–1345 1365–1375 1395–1405 1425–1435 1455–1465 1485–1495 1515–1525 1545–1555 1575–1585 1605–1615 1635–1645 1665–1675 1695–1705 1725–1735 1755–1765 1785–1795 1815–1825 1845–1855 1875–1885 1905–1915 1935–1945 1965–1975 1995–2005 2025–2035 2055–2065 520–530 550–560 580–590 610–615 640–650 670–680 700–710 730–740 760–770 790–800 820–830 850–860 880–890 910–920 940–950 970–980 1000–1010 1030–1040 1060–1070 1090–1100 1120–1130 1150–1160 1180–1190 1210–1220 1240–1250 1270–1280 1300–1310 1330–1340 1360–1370 1390–1400 1420–1430 1450–1460 1480–1490 1510–1520 1540–1550 1570–1580 1600–1610 1630–1640 1660–1665 1700–1710 1720–1730 1750–1760 1780–1790 1810–1820 1840–1850 1870–1880 1900–1910 1930–1940 1960–1965 1990–2000 2020–2030 2050–2060 depositional environment depositional environment n o n -m ar in e c o as ta l – m ar gi n al m ar in e in n er n er it ic o u te r n er it ic o p en o ce an n o n -m ar in e c o as t a l – m ar gi n al m ar in e in n er n er it ic o u te r n er it ic o p en o ce an 600 (m) 700 800 900 1000 1100 1200 1300 1400 1500 1600 1700 1800 1900 2000 thanetian selandian danian la te maastrichtian late campanian i late campanian ii e. campanian ea rl y 2085–2095 2115–2125 2145–2155 2175–2185 2205–2215 2235–2245 2265–2275 2295–2305 2325–2335 2350–2360 2380–2390 2410–2420 2440–2450 2470–2480 2500–2510 2530–2540 2560–2570 2590–2600 2620–2630 2650–2660 2670–2680 2700–2710 2730–2740 2760–2770 2790–2800 2820–2830 2850–2860 2880–2890 2910–2920 2940–2950 2970–2980 3000–3010 3030–3040 3060–3070 3090–3100 3120–3130 3150–3160 3180–3190 3210–3220 3240–3250 3270–3280 3300–3310 3330–3340 3360–3370 3390–3400 3420–3430 3450–3460 3480–3490 3510–3520 3540–3550 3570–3580 3595–3605 2080–2090 2110–2120 2135–2145 2170–2180 2200–2210 2240–2250 2260–2270 2290–2300 2320–2330 2360–2370 2560–2570 2680–2690 2705–2715 2740–2750 2770–2780 2800–2810 2830–2840 2860–2870 2890–2900 2920–2930 2950–2960 2980–2990 3010–3020 3040–3050 3070–3080 3100–3110 3130–3140 3160–3170 3190–3200 3220–3230 3250–3260 3280–3290 3310–3320 3340–3350 3370–3380 3410–3420 3430–3440 3460–3470 3490–3500 3520–3530 3550–3560 3580–3590 2100 2200 2300 2400 2500 2600 2700 2800 2900 3000 3100 3200 3300 3400 3500 3600 age suite 1 suite 2 period/ epoch sample depth (m) age suite 1 suite 2 period/ epoch sample depth (m) basal ypresian m io c e n e o l ig o c e n e e o c e n e e o c e n e p a l e o c e n e c r e t a c e o u s l l e e e l l e m l e-m appendix 2.4. stratigraphy and palaeoenvironmental interpretation of the gilbert f-53 well. kelly bushing height: 12.1 m above sea level. water depth: 183 m. total depth: 3608.0 m. bulletin37.qxp_bulletin 37 26/01/17 14.00 side 71 72 800 900 1000 1100 1200 1300 1400 1500 1600 1700 1800 1900 2000 2100 2200 2300 2400 2500 miocene chattian rupelian priabonian bartonian lutetian 800–810 890–900 920–930 950–960 980–990 1060–1070 1090–1100 1120–1130 1150–1160 1180–1190 1230–1240 1260–1270 1290–1300 1330–1340 1360–1370 1390–1400 1420–1430 1460–1470 1490–1500 1520–1530 1550–1560 1580–1590 1610–1620 1640–1650 1660–1670 1720–1730 1750–1760 1780–1790 1840–1850 1870–1880 1900–1910 1930–1940 1970–1980 2000–2010 2030–2040 2060–2070 2090–2100 2120–2130 2150–2160 2180–2190 2210–2220 2240–2250 2270–2280 2300–2310 2330–2340 2360–2370 2390–2400 2420–2430 2460–2470 2490–2500 2520–2530 2550–2560 1040–1050 1110–1120 1220–1230 1465–1475 1510–1520 1650–1660 1690–1700 1805–1815 1895–1905 2080–2090 2170–2180 2200–2210 2230–2240 2215–2225 2260–2270 2365–2375 2391–2401 2435–2445 2470–2480 2510–2520 2551–2561 1816.0 2435.6 2587.0 2600 2610–2620 2700 2800 2900 3000 3100 3200 3300 3400 3500 3600 3700 3800 3900 4000 4100 4200 4300 4400 4500 ypresian ypresian thanetian selandian 2670–2680 2700–2710 2730–2740 2790–2800 2820–2830 2850–2860 2880–2890 2895–2905 2920–2930 2940–2950 2970–2980 3000–3010 3030–3040 3060–3070 3090–3100 3120–3130 3140–3150 3150–3160 3180–3190 3210–3220 3270–3280 3310–3320 3350–3360 3410–3420 3450–3460 3510–3520 3570–3580 3610–3620 3650–3660 3730–3740 3770–3780 3830–3840 3870–3880 3930–3940 3970–3980 4010–4020 4030–4040 4050–4060 4070–4080 4090–4100 4110–4120 4130–4140 4150–4160 4170–4180 4190–4200 4210–4220 4230–4240 4250–4260 4270–4280 4290–4300 4310–4320 4330–4340 4350–4360 4370–4380 4390–4440 4410–4420 4430–4440 4450–4460 4470–4480 4490–4500 4510–4520 4530–4540 4550–4560 4340–4350 4353–4363 4380–4390 4400–4410 4432–4442 4420–4430 4440–4450 4460–4470 4480–4490 4500–4510 4538–4548 4540–4550 4560–45664565–4566 4566–4566 2655–2665 2690–2700 2740–2750 2770–2780 2870–2880 2920–2930 2930–2940 2931–2941 2960–2970 2985–2995 3045–3055 3091–3101 3121–3131 3139–3149 3141–3151 3160–3170 3181–3191 3200–3210 3211–3221 3225–32353230–3240 3245–32553251–3261 3265–3275 3271–3281 3300–3310 3360–3370 3320–3330 3400–3410 3480–3490 3500–3510 3520–3530 3600–3610 3640–3650 3700–3710 3630–3640 3680–3690 3720–3730 3760–3770 3840–3850 3795–3805 3800–3810 3880–3890 3920–3930 3900–3910 3960–3970 4000–4010 4040–4050 4080–4090 4120–4130 4180–4190 4200–4210 4220–4230 4260–4270 4300–4310 2766.0 2918.0 3245.8 3250.0 3258.0 3546.0 4100–4110 4001–4011 2640–2650 n o n -m ar in e c o as ta l – m ar gi n al m ar in e in n er n er it ic o u te r n er it ic o p en o ce an depositional environment depositional environment n o n -m ar in e c o as ta l – m ar gi n al m ar in e in n er n er it ic o u te r n er it ic o p en o ce an age period/ epoch sample depth (m) m io c e n e o l ig o c e n e e o c e n e e o c e n e e l l m e e p a l e o c e n e l suite 1 suite 2 age period/ epoch sample depth (m) suite 1 suite 2 swc appendix 2.5. stratigraphy and palaeoenvironmental interpretation of the hekja o-71 well. kelly bushing height: 12.5 m above sea level. water depth: 350.8 m. total depth: 4566.0 m. bulletin37.qxp_bulletin 37 26/01/17 14.00 side 72 73 1500 (m) 1600 1700 1800 1900 2000 2100 2200 2300 2400 2500 2600 2700 rupelian priabonian bartonian lutetian thanetian 2050–2060 2110–2120 2190–2200 2290–2300 1490–1500 1510–1520 1530–1540 1520–1530 1550–1560 1565–1575 1570–1580 1575–1585 1600–1610 1620–1630 1610–1620 1640–1650 1630–1640 1660–1670 1670–1680 1690–1700 1700–1710 1710–1720 1740–1750 1760–1770 1750–1760 1780–1790 1770–1780 1800–1810 1820–1830 1840–1850 1830–1840 1858–1868 1860–1870 1890–1900 1950–1960 1920–1930 1990–2000 2020–2030 2030–2040 2070–2080 2150–2160 2200–2210 2230–2240 2250–2260 2310–2320 2270–2280 2320–2330 2350–2360 2380–2390 2390–2400 2410–2420 2440–2450 2470–2480 2430–2440 2500–2510 2510–2520 2530–2540 2550–2560 2580–2590 2590–2600 2600–2610 2620–2630 2650–2660 2670–2680 2630–2640 2680–2690 2690–2700 2710–2720 2800 (m) 2900 3000 3100 3200 3300 3400 3500 3600 3700 3800 3900 thanetian selandian danian 2740–2750 2750–2760 2770–2780 2800–2810 2790–2800 2830–2840 2840–2850 2870–2880 2890–2900 2920–2930 2917–2927 2940–2950 2960–2970 2950–2960 3000–3010 2990–3000 3030–3040 3060–3070 3090–3100 3070–3080 3120–3130 3110–3120 3150–3160 3180–3190 3190–3200 3200–3210 3230–3240 3270–3280 3300–3310 3320–3330 3330–3340 3350–3360 3360–3370 3390–3400 3420–3430 3450–3460 3470–3480 3430–3440 3480–3490 3510–3520 3540–3550 3550–3560 3560–3570 3590–3600 3620–3630 3600–3610 3650–3660 3680–3690 3670–3680 3710–3720 3740–3750 3770–3780 3790–3800 3750–3760 3810–3820 3830–3840 3860–3870 3890–3900 3870–3880 3910–3920 3900–3910 3940–3950 3920–3930 3970–3980 3980–3990 3950–3960 3960–3970 3988–3998 depositional environment depositional environment n o n -m ar in e c o as ta l – m ar gi n al m ar in e in n er n er it ic o u te r n er it ic o p en o ce an age period/ epoch age sample depth (m) period/ epoch sample depth (m) n o n -m ar in e c o as ta l – m ar gi n al m ar in e in n er n er it ic o u te r n er it ic o p en o ce an ypresian e o c e n e p a l e o c e n e p a l e o c e n e o l i e l m e e l l appendix 2.6. stratigraphy and palaeoenvironmental interpretation of the gjoa g-37 well. kelly bushing height: 24.0 m above sea level. water depth: 1000 m. total depth: 3998 m. bulletin37.qxp_bulletin 37 26/01/17 14.00 side 73 74 appendices 3.1–3.19 are located in pockets attached to the inside covers palynological event summary charts for the 19 wells used in this study, in order from south to north (for well locations, see fig. 1). although the information included varies from chart to chart, wireline logs, a simplified lithology log, lithostratigraphic units, palynological events and chronostratigraphic subdivisions are included for all wells. charts of wells analysed by henrik nøhr-hansen (hellefisk-1, ikermiut-1, kangâmiut-1, nukik-1, nukik-2, qulleq-1, north leif i-05, ogmund e-72, skolp e-07, hekja o-71, ralegh n-18 and gjoa g-37) include, in addition to palynological events, a biostratigraphic zonation for the aptian–albian to priabonian of the west greenland and labrador margins of the labrador sea; this zonation was published by nøhr-hansen in sønderholm et al. (2003b), nøhr-hansen et al. (2000, 2002) and nøhr-hansen (2003, 2004a, b). some of the well charts listed above also show dinocyst species diversity, azolla abundances and palaeoenvironmental interpretations for the palaeogene (after rasmussen et al. 2003). plots generated by graham l. williams and robert a. fensome (bjarni o-82, gilbert f-53, karlsefni a-13, roberval k-92, rut h-11, snorri j-90 and south labrador n-79) show events and ages only; a biozonation was not developed. kb: kelly bushing. appendix 3 bulletin37.qxp_bulletin 37 26/01/17 14.00 side 74 75 bulletin37.qxp_bulletin 37 26/01/17 14.00 side 75 76 de nationale geologiske undersøgelser for danmark og grønland (geus) geological survey of denmark and greenland øster voldgade 10, dk-1350 copenhagen k denmark the series geological survey of denmark and greenland bulletin started in 2003 and replaced the two former bulletin series of the survey, viz. geology of greenland survey bulletin and geology of denmark survey bulletin. the twenty-one volumes published since 1997 in those two series are listed on the following pages. the present series, together with geological survey of denmark and greenland map series, now form the peer-reviewed scientific series of the survey. geological survey of denmark and greenland bulletin 1 the jurassic of denmark and greenland, 948 pp. (28 articles), 2003. edited by j.r. ineson & f. surlyk. 500.00 2 fish otoliths from the paleocene of denmark, 94 pp., 2003. by w. schwarzhans. 100.00 3 late quaternary environmental changes recorded in the danish marine molluscan faunas, 268 pp., 2004. by k.s. petersen. 200.00 4 review of survey activities 2003, 100 pp. (24 articles), 2004. edited by m. sønderholm & a.k. higgins. 180.00 5 the jurassic of north-east greenland, 112 pp. (7 articles), 2004. edited by l. stemmerik & s. stouge. 160.00 6 east greenland caledonides: stratigraphy, structure and geochronology, 93 pp. (6 articles), 2004. edited by a.k. higgins & f. kalsbeek. 160.00 7 review of survey activities 2004, 80 pp. (19 articles), 2005. edited by m. sønderholm & a.k. higgins. 180.00 8 structural analysis of the rubjerg knude glaciotectonic complex, vendsyssel, northern denmark, 192 pp., 2005. by s.a.s. pedersen. 300.00 9 scientific results from the deepened lopra-1 borehole, faroe islands, 156 pp. (11 articles), 2006. edited by j.a. chalmers & r. waagstein. 240.00 10 review of survey activities 2005, 68 pp. (15 articles), 2006. edited by m. sønderholm & a.k. higgins. 180.00 11 precambrian crustal evolution and cretaceous–palaeogene faulting in west greenland, 204 pp. (12 articles), 2006. edited by a.a. garde & f. kalsbeek. 240.00 12 lithostratigraphy of the palaeogene – lower neogene succession of the danish north sea, 77 pp., 2007. by p. schiøler, j. andsbjerg, o.r. clausen, g. dam, k. dybkjær, l. hamberg, c. heilmann-clausen, e.p. johannessen, l.e. kristensen, i. prince & j.a. rasmussen. 240.00 13 review of survey activities 2006, 76 pp. (17 articles), 2007. edited by m. sønderholm & a.k. higgins. 180.00 14 quaternary glaciation history and glaciology of jakobshavn isbræ and the disko bugt region, west greenland: a review, 78 pp., 2007. by a. weidick & o. bennike. 200.00 15 review of survey activities 2007, 96 pp. (22 articles), 2008. edited by o. bennike & a.k. higgins. 200.00 16 evaluation of the quality, thermal maturity and distribution of potential source rocks in the danish part of the norwegian–danish basin, 66 pp., 2008. by h.i. petersen, l.h. nielsen, j.a. bojesen-koefoed, a. mathiesen, l. kristensen & f. dalhoff. 200.00 17 review of survey activities 2008, 84 pp. (19 articles), 2009. edited by o. bennike, a.a. garde & w.s. watt. 200.00 18 greenland from archaean to quaternary. descriptive text to the 1995 geological map of greenland, 1:2 500 000. 2nd edition, 126 pp., 2009. by n. henriksen, a.k. higgins, f. kalsbeek & t.c.r. pulvertaft. 280.00 19 lithostratigraphy of the cretaceous–paleocene nuussuaq group, nuussuaq basin, west greenland, 171 pp., 2009. by g. dam, g.k. pedersen, m. sønderholm, h.h. midtgaard, l.m. larsen, h. nøhr-hansen & a.k. pedersen. 300.00 20 review of survey activities 2009, 106 pp. (23 articles), 2010. edited by o. bennike, a.a. garde & w.s. watt. 220.00 21 exploration history and place names of northern east greenland, 368 pp., 2010. by a.k. higgins. 200.00 22 lithostratigraphy of the upper oligocene – miocene succession of denmark, 92 pp., 2010. by e.s. rasmussen, k. dybkjær & s. piasecki. 240.00 23 review of survey activities 2010, 84 pp. (19 articles), 2011. edited by o. bennike, a.a. garde & w.s. watt. 200.00 bulletin37.qxp_bulletin 37 26/01/17 14.00 side 76 77 24 the east greenland rifted volcanic margin, 96 pp., 2011. by c.k. brooks. 200.00 25 upper cretaceous chalk facies and depositional history recorded in the mona-1 core, mona ridge, danish north sea. 2011. by k. anderskouv & f. surlyk. 200.00 26 review of survey activities 2011, 88 pp. (21 articles), 2012. edited by o. bennike, a.a. garde & w.s. watt. 200.00 27 neoglacial and historical glacier changes around kangersuneq fjord in southern west greenland, 68 pp., 2012. by a. weidick, o. bennike, m. citterio & n. nørgaard-pedersen. 200.00 28 review of survey activities 2012, 76 pp. (17 articles), 2013. edited by o. bennike, a.a. garde & w.s. watt. 200.00 29 tectono-magmatic evolution of the younger gardar southern rift, south greenland, 124 pp., 2013. by b.g.j. upton. 240.00 30 stratigraphic landscape analysis, thermochronology and the episodic development of elevated, passive continental margins, 150 pp., 2014. by paul f. green, karna lidmar-bergström, peter japsen, johan m. bonow and james a. chalmers. 250.00 31 review of survey activities 2013, 98 pp., 2014. edited by o. bennike, a.a. garde & w.s. watt 200.00 32 a catalogue of danian gastropods from the baunekule facies, faxe formation, denmark, 117 pp., 2014 by b.w. lauridsen & k.i. schnetler. 240.00 33 review of survey activities 2014, 88 pp. (20 articles), 2015. edited by o. bennike, a.a. garde & w.s. watt. 200.00 34 the ammonites of the middle jurassic cranocephalites beds of east greenland, 145 pp., 2015. by j.h. callomon, p. alsen & f. surlyk. 250.00 35 review of survey activities 2015, 106 pp. (24 articles), 2015. edited by a.a garde, o. bennike, k. thrane & w.s. watt. 200.00 36 cretaceous and cenozoic dinoflagellate cysts and other palynomorphs from the western and eastern margins of the labrador–baffin seaway, 143 pp., 2016. by r.a. fensome, h. nøhr-hansen & g.l. williams. 250.00 37 biostratigraphic correlation of the western and eastern margins of the labrador–baffin seaway and implications for the regional geology, 75 pp., 2016. by h. nøhr-hansen, g.l. williams & r.a. fensome. 200.00 geological survey of denmark and greenland map series 1 explanatory notes to the geological map of greenland, 1:500 000, humboldt gletscher, sheet 6, 48 pp. + map, 2004. by p.r. dawes. 280.00 2 explanatory notes to the geological map of greenland, 1:500 000, thule, sheet 5 (1991), 97 pp. + map, 2006. by p.r. dawes. 300.00 3 explanatory notes to the geological map of greenland, 1:100 000, ussuit 67 v.2 nord, 40 pp. + map, 2007. by j.a.m. van gool & m. marker. 280.00 4 descriptive text to the geological map of greenland, 1:500 000, dove bugt, sheet 10, 32 pp. + map, 2009. by n. henriksen & a.k. higgins. 240.00 5 descriptive text to the geological map of greenland, 1:100 000, kangaatsiaq 68 v.1 syd and ikamiut 68 v.1 nord, 41 pp. + 2 maps, 2010. by a.a. garde & j.a. hollis. 280.00 geology of greenland survey bulletin (173–191; discontinued) 173 cambrian shelf stratigraphy of north greenland, 120 pp., 1997. by j.r. ineson & j.s. peel 250.00 174 the proterozoic thule supergroup, greenland and canada: history, lithostratigraphy and development, 150 pp., 1997. by p.r. dawes 300.00 175 stratigraphy of the neill klinter group; a lower – lower middle jurassic tidal embayment succession, jameson land, east greenland, 80 pp., 1998. by g. dam & f. surlyk. 250.00 176 review of greenland activities 1996, 112 pp. (18 articles), 1997. edited by a.k. higgins & j.r. ineson. 200.00 177 accretion and evolution of an archaean high-grade grey gneiss – amphibolite complex: the fiskefjord area, southern west greenland, 115 pp., 1997. by a.a. garde. 200.00 178 lithostratigraphy, sedimentary evolution and sequence stratigraphy of the upper proterozoic lyell land group (eleonore bay supergroup) of east and north-east greenland, 60 pp., 1997. by h. tirsgaard & m. sønderholm. 200.00 bulletin37.qxp_bulletin 37 26/01/17 14.00 side 77 << /ascii85encodepages false /allowtransparency true /autopositionepsfiles true /autorotatepages /none /binding /left /calgrayprofile (dot gain 20%) /calrgbprofile (apple rgb) /calcmykprofile (u.s. web coated \050swop\051 v2) /srgbprofile (srgb iec61966-2.1) /cannotembedfontpolicy /warning /compatibilitylevel 1.6 /compressobjects /tags /compresspages true /convertimagestoindexed true /passthroughjpegimages true 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/mobilecompatible 0 /namespace [ (adobe) (golive) (8.0) ] /openzoomtohtmlfontsize false /pageorientation /portrait /removebackground false /shrinkcontent true /treatcolorsas /mainmonitorcolors /useembeddedprofiles false /usehtmltitleasmetadata true >> ] >> setdistillerparams << /hwresolution [2400 2400] /pagesize [612.000 792.000] >> setpagedevice untitled 120 121 from the 30 m décollement level. the hanging-wall flat of rf06 rests on the upper flat of rf05, where no deposits of the rubjerg knude formation have been recognised. thus the two thrust sheets occur as a block of lønstrup klint formation 60 m thick, separated in the middle by a thrust fault. the l/r-unconformity in the rf06 sheet is located at about 10–15 m above sea level. thus, the rf06 sheet was faulted up on an intermediate flat above the trailing edge of rf05, probably while both were transported along the lower footwall flat on top of the lower trailing duplex segment of the grønne rende section. sedimentary units in the rubjerg knude fyr section, the lower stratigraphic levels of the lønstrup klint formation are exposed, although they are commonly deformed either by mudmobilisation or thrust-fault shearing. the rubjerg knude formation above is poorly exposed, partly due to sand scree derived in part from the formation itself, and partly from the up to 50 m high sand dunes above the cliff. no further description of the formations is given here. facing page – upper: fig. 99. normal faults displacing the top of the rf03 thrust sheet in the rubjerg knude fyr section. note the footwall syncline folded below the hanging-wall ramp at the top of the cliff section. photograph: june 1984. facing page – lower: fig. 100. normal fault (nf) displacing the tip of the rf04 thrust sheet, which prior to normal faulting was thrusted along the hanging-wall flat of rf04 (rf04hwf). photograph: july 1994. fig. 101. thrust-fault brecciation related to the lower hanging-wall ramp and flat in the rubjerg knude fyr section. the brecciation fabric that is typical of the fine-grained sand turbidites and laminated clayey muds (compare with primary sedimentary features in figs 22, 23) was produced by low-angle anastomosing shearing. photograph: september 1985. 122 structures two types of deformation alter the primary sedimentary architecture of the rubjerg knude fyr section. the first type is anastomosing thrust-fault brecciation related to the zone above the hanging-wall ramp with fine-grained turbidites. the second type is the mudmobilisation and mesoscopic-scale polydiapirism, which is common in the clay-rich lower part of the thrust sheets. these deformation types have to be considered in the evaluation of the balance calculation. only the anastomosing thrust-fault brecciation will be further described in this section; descriptions of diapirism are given under the kramrende, brede rende, sandrende and moserende sections. anastomosing thrust-fault brecciation the most significant mesoscopic-scale structure recognised in the rubjerg knude fyr section is related to thrust-fault brecciation in the lower part of the lønstrup klint formation. the brecciation fabric that is typical of the fine-grained turbidites and laminated clayey muds was created by low-angle anastomosing shearing. the shear surfaces and thrust-fault displacements are located in the clay-rich laminae, whereas the segments bounded by the anastomosing fractures consist of silty mud lithologies (fig. 101). the brecciation extends from the lower hanging-wall ramp-andflat up to 10 m above the sole of the sheets. the anastomosing thrust faulting indicates that significant dif ferential movements developed in the subsurface during thrust-fault propagation, which illustrates the nature of the displacement-related shearing and which may account for some of the volume problems arising from construction of the balanced crosssection. interpretation of structural development the compression (shortening) in the rubjerg knude fyr section is about 48%, calculated from the measured length of the section of 270 m (= l 1 ) and the balanced length of about 520 m (= l 0 ). this implies that the thrust sheets in the section have been displaced upwards at the ramp originally situated at the trailing end of rf04 and transported to the position of the footwall ramp of gr13. the lower flat is still the footwall flat on top of the lower duplex segment of the grønne rende section. thus, a considerable amount of lateral translation is apparent in the rubjerg knude fyr section. the appearance of the normal fault with down-faulted noses of the rf03 and rf04 sheets indicates that differential movements, including duplex development of the gr01–gr06 lower segments, may have occurred to create the foreland-dipping features in the rf02– rf03 piggyback basin. the interesting structure in rf05 is the lower hanging-wall ramp, which has to correspond to a footwall ramp at the trailing end of rf04. the footwall ramping probably also included propagation along a footwall ramp of a lower duplex segment of rf04 (rf04u). the present orientation of the rf05 lower hangingwall ramp is more or less vertical, indicating three steps of ramping. the final tilting was due to the ramping of the upper footwall ramp of gr13, the middle phase of ramping was up along the footwall ramp of the piggyback basin in the normal fault displaced rf04, and the initial ramping was probably a complex propagation over several smaller ramp-steps that separated rf06 from rf05. the uppermost part of rf06 shows a marked topography, indicating that at an early phase of deformation it was elevated up to a level of erosion, before subsequent sedimentation. this sedimentation was probably of relatively short duration before over-thrusting of the stortorn section trapped the piggyback basin. stortorn section stortorn is the name of the very steep and muddy cliff in the central part of the rubjerg knude cliff section. on old drawings of the beach and the coastal cliff, stortorn is depicted as a steep, wild looking castlelike cliff in the distant horizon, emphasising the romantic scenery of this remote place (e.g. engraving by c. neumann 1884, reproduced in vendsyssel nu og da, 1981). in recent times, the cliff has also been the location of major landslides, which in some cases travelled more than 100 m out into the sea. in general, the sea reaches close up to the vertical cliff, and due to the muddy and slippery cliff surfaces and the clay pavement in the zone of breakers, it is the most difficult place to pass along the coast. in the stortorn section, the deepest level of thrusting occurs where the décollement zone is located at a depth of 40 m stratigraphically below the reference surface of the l/r-unconformity, which is about 45 m 123 below sea level. from this deep level, the thrust sheets were elevated up to the exposed position in the cliff section. coinciding with this, the thrust sheets contain the deepest levels of the stratigraphy, and beds of marine and glaciomarine clay can be identified by the occurrences of arctic marine fossils. moreover, the stortorn section contains the key features for understanding the structural and dynamic problems of the adjacent rubjerg knude fyr and grønne rende sections. the key features are flat-lying duplex complexes formed by ramping up of the relatively long, lower thrust-sheet segments onto a high flat level. this is reflected in an elevation of the l/r-unconformity up to a height of c. 40 m above sea level in the cliff section. tectonic architecture the stortorn section is divided into ten thrust sheets annotated st01–st10. the southern boundary of the section is the frontal hanging-wall flat of st01, which coincides with the footwall ramp of the rf06 thrust sheet in the rubjerg knude fyr section. the northern boundary is the trailing footwall thrust of st10 along which the frontal hanging-wall ramp of the moserende section was thrust. the southernmost three thrust sheets form a separate group of high-level thrust sheets. the l/r unconformity is here situated at an elevation of 35–40 m above sea level. the central part of the section is formed by a series of thick thrust sheets of clayey, partly mobilised, mud, which are situated above hidden duplexes in the subsurface. this complex extends about 150 m along the cliff section at stortorn. in the northern part of the section, upright mud diapir-dominated thrust sheets occur with complexly developed piggyback basins. st01 thrust sheet the southernmost thrust sheet in the stortorn section (st01) is wedge-shaped, c. 160 m long, with an initial 25° dip of the frontal hanging-wall ramp. the footwall ramp (fr06 trailing edge) dips at about 45°, which creates a problem in the balancing. it is obvious that a splint (or horse) corresponding to a triangle with an acute angle of 20° must be hidden somewhere in the deeper structure. the trailing end of st01 is 30 m thick. however, the lønstrup klint formation is deeply eroded in the central part of st01, which truncates the l/r-unconformity. st02 thrust sheet the st02 thrust sheet was also elevated to a height of 35–40 m in the cliff section. the hanging-wall ramp is vertical and forms a right-angle with the horizontal bedding in the piggyback basin of st01. this indicates that final up-thrusting of the frontal hanging-wall ramp took place in an upright position during sedimentation on the back of st01. the st02 thrust fault was rotated at least three times before propagation up along the footwall ramp of st01. the initial dip of this ramp was relatively steep, about 40°, as indicated by the angle between the thrust fault and the bedding in the lønstrup klint formation of st01. furthermore, the problem related to the change in ramp angle recurs. in the balancing of the thrust structure, a splint volume must be calculated for, corresponding to the triangle created by the initial thrust angle and the final steeply inclined thrust fault. st03 thrust sheet the st03 thrust sheet is not elevated as much as st01 and st02; the l/r-unconformity is only situated at about 30 m above sea level. the lønstrup klint formation was deeply eroded before sedimentation in the piggyback basin was initiated, probably due to marked relief during ramp propagation. st04 thrust sheet the lowest position of the l/r-unconformity in the st04 thrust sheet is about 13–15 m above sea level, which demonstrates a shallower level of ramping than in st01– st03. the 45° steeply dipping thrust fault between st03 and st04 roots down to the décollement zone and propagated up along the footwall ramps of the subsurface duplex segments beyond st01–st03. the dif ference between initial and final angle of ramping also created a balancing problem for st04. the frontal part of the hanging-wall ramp of st04 only had a dip of 18–20°, whereas the footwall ramp now dips at c. 45°. therefore a splint (with an area of 630 m2 in the cross-section) is envisaged in the subsurface. the st04 ramping over this splint is interpreted as the rea124 son for the elevation of the l/r-unconformity up to c. 15 m above sea level. st05 thrust sheet the st05 thrust sheet is one of the most important structures, not only in the stortorn section but also in the rubjerg knude glaciotectonic complex as a whole. it involved thrusting of the deepest décollement level, which introduced a complex framework due to the large number of duplex segments involved. the st05 thrust sheet was thrust up along the footwall ramp of st04. the dip of the thrust fault increases from 35° at the beach level to 70° at the top of the cliff section. the tip of the thrust sheet was finally displaced horizontally over the top of the piggyback basin of st04. the thrust displacement along the hanging-wall thrust fault is estimated to be in the order of 90 m. the lønstrup klint formation of st05 is dominated by mud diapirism and polydiapiric structures up to 5 m in vertical scale. in the middle part of the thrust sheet, where the thickness of the initially wedgeshaped frontal part was 20 m, a deeply eroded trough was formed. the l/r-unconformity on the northern flank of this erosional depression is situated nearly 50 m above sea level. this is about the highest elevation of the reference surface, and was caused by thrust duplication of the thrust sheet in the subsurface duplex complex. st06 thrust sheet the st06 thrust sheet is poorly exposed and mudmobilisation and internal diapirism obscure primary structures. the thickness of the thrust sheet is up to 40 m, measured from the frontal hanging-wall thrust up to a small pocket of rubjerg knude formation sand that forms the remnant of a piggyback basin. the frontal thrust is drawn with some uncertainty, because a large part of it is penetrated by diapirism intruding from the back of st05. a minimum displacement of 40 m is inferred, which is incorporated in the modelling of the balanced cross-section. this implies a rather complex structural assemblage of the subsurface lower duplex segments of st05, st06 and st07. st07 thrust sheet the st07 thrust sheet is nearly vertically orientated with a frontal ramp rising from 70° to vertical, along which a displacement of 28 m is estimated to have occurred. the l/r-unconformity surface is also steeply dipping, and is even overturned at the top. the lower part of the thrust sheet is mainly covered by scree, but it is possible to trace the line of the unconformity down to the level of the beach in the crosssection. this implies that the thrust sheet has not been thrust up to be displaced along an intermediate ramp but is only tilted due to the main ramping, first along the thrust fault of st06 and finally on its own hanging-wall ramp. the ramping is also reflected in the deposition in the piggyback basin where three superposed angular discordances are recognised. the deposits in the piggyback basin are c. 15 m thick; at the base of the succession, the r-onlap starts with an angle of 45° and terminates with a 90° angle, indicating deposition in the basin while the thrust sheet was vertically orientated. in the uppermost bed, minor slump folds are present, indicating the effect of the st08 thrust nose approaching from the north. st08 thrust sheet a double ramp synclinal structure, similar to the one occurring in the piggyback basin of st05, is recognised in thrust sheet st08. the l/r-unconformity incises through the lønstrup klint formation and down into the thrust-fault surface of the hanging-wall ramp. the lateral distance between the ramps is only about 25 m and the basin is less than 10 m deep. the sediments in this piggyback basin show large-scale trough cross-bedding accentuated by synclinal folding. the rubjerg knude formation covers a feature that represents the erosional remnants of a detachment anticline on the northern limb of the basin. on the north side of this structure, the initial stratification above the l/r-unconformity shows clear r-onlap, corresponding to the inclination parallel to the tilt of the lower ramp. the unconformity is elevated up to 20 m above sea level, indicating that the st08 thrust sheet was lifted up by at least two subsurface duplex segments. these hidden segments are annotated st08u1 and st08u2. the displacement along the footwall ramp of st08 is estimated at 78 m, mainly along the thrust fault dipping at 30°. 125 st09 thrust sheet the st09 thrust sheet is a c. 30 m thick sheet bounded by a 60° dipping hanging-wall flat thrust up onto the footwall ramp of st08 back and the 60° dipping footwall thrust fault of st10, which truncates the irregular structures in the upper part of the thrust sheet. the displacement along the hanging-wall ramp is estimated at c. 60 m. although the lønstrup klint formation in st09 is characterised by internal diapirism, the features of a hanging-wall anticline can be recognised at the top of the cliff section. during translation of two or more footwall ramps, an irregular synform formed and created the depocentre of a piggyback basin. the l/r-unconformity is here elevated to 5–10 m above sea level, corresponding to propagation up onto the st08u2 duplex segment. st10 thrust sheet the st10 thrust sheet is about the same size as the st09 sheet, and also has steeply dipping bounding thrust faults. the most remarkable structure in the st10 thrust sheet is the structural complexity of the piggyback basin. the l/r-unconformity forms an isoclinal recumbent syncline, with the upper limb formed by the mud-mobilised lønstrup klint formation, and above this a minor synclinal trough appears. this structure is best described as a detachment anticline, which developed into a diapir with a reverse fault displacing the northern limb of the structure into a mushroomshaped structure, similar to the diapir in the sandrende section. at a late stage of thrusting, the piggyback basin was rotated 60° and the diapir-developed detachment anticline collapsed into the recumbent strucfig. 102. isoclinal upright anticline formed in the lower part of the lønstrup klint formation in the frontal part of the stortorn section. the right limb of the anticline constitutes an imbricate duplex formed by connecting thrust-fault splays (white dot-and-dash lines). the fold is interpreted as a hanging-wall anticline developed during fault propagation and successive imbricate stacking (compare with fig. 59). photograph: august 2001. 126 ture. the l/r-unconformity is elevated up to 15–18 m above sea level indicating a ramping of the st08u duplex segment as well as the trailing segment of st09. an estimated displacement of 73 m along the hanging-wall ramp of st10 still leaves some subsurface segments to be balanced in the structure below the moserende section to the north. sedimentary units the most important sedimentary feature in the stortorn section is the exposure of the stortorn formation, which is the lowermost stratigraphic level involved in the rubjerg knude glaciotectonic complex. the stortorn formation is located at the lower hanging-wall ramp of the st05 thrust sheet, where it forms a duplex segment about 3–5 m thick at the base of the cliff. to date, it has not been possible to measure a sedimentological log of the formation at this locality, partly because the formation is strongly sheared by anastomosing fractures, and partly because landslide activity precludes more detailed stratigraphic description. structures the contact between mobilised, intrusive mud and stratified mud has been observed in many places, but this characteristic is better illustrated in the moserende section (see below). anastomosing thrust faults and tectonic breccias occur commonly in the st05 and st06 thrust sheets. however, due to the difficult field conditions, detailed investigations have not been carried out. in the dark clayey mud of thrust sheet st01, an upright nearly isoclinal anticline has been observed (fig. 102). this fold is considered to represent a hanging-wall anticline that was subjected to an advanced stage of deformation during ramp propagation. this stage compares well with the model of duplex formation described by mitra & sussman (1997), in which the growth of imbricates derived from successive connecting splays results in steepening of antiformal stacks formed by fault-propagation folding of duplexes. successive growth of duplex elements corresponds well with the interpretation presented below. interpretation of structural development the cross-section (plate 2b) provides a model for the structures below the frontal part of the stortorn section, which requires four duplex segments forming a duplex complex, on top of which thrust sheets st01, st02 and st03 have been thrust along the footwall flat. the structural interpretation of a duplex stacking of subsurface segments below the frontal part of the stortorn section is based on two lines of evidence: (1) the missing balance of the lower segments related to the grønne rende section, and (2) the high elevated position of the l/r-unconformity in this part of the section. thus the lower duplex complex in the frontal part of the section is interpreted to represent stacking of the trailing lower segment of gr01, although an alternative differential duplex-segment displacement is also possible. from geometric considerations, it is evident that the 20 m and 30 m décollement levels must have been pervasive throughout the proximal part of the thrust structure. thus the duplex in the stortorn section consists of segments c. 10 m thick. mobilisation and the internal polydiapirism have obscured the boundaries of these segments, which are the lower and intermediate footwall and hanging-wall flats respectively. however, in a model for reconstruction, these volumes are regarded as solid thrust sheet, i.e. the duplex segments (represented by annotated areas in plate 2). in the description and solution of the structural problem related to differential thrust faulting of the lower duplex segments, three main types of fault-bend-folded segments are distinguished (fig. 103). fig. 103. schematic illustration of the three types of fault-bend folding of duplex segments. type 1 is referred to as an l-structure, type 2 as an s-structure and type 3 as a g-structure (g chosen due to similarity with the greek capital letter gamma (γ)). the footwall ramp dips at about 45°, the shortening between the underlying footwall ramp and the overlying hanging-wall ramp is 43%, the initial length (l 0 ) of the thrust sheet is c. 100 m, compared to a thrust-sheet thickness of 40 m, and a thickness of 10 m for the individual duplex segments. 127 1. a duplex segment with one part of the segment resting on the lower flat and the other parallel with a ramp (l-structure). 2. a duplex segment which has the trailing part parallel with the lower flat, the intermediate part parallel with a ramp, and the frontal part parallel with the upper flat, thus giving the shape of a s. 3. a duplex segment with the trailing part of the segment located parallel with the ramp and the frontal part parallel with the upper flat (γ-structure). the stacking of the duplex below the st01, st02 and st03 thrust sheets started when the lower segments were thrust up into the first type duplex bend during ramping towards the footwall ramp of rf06 (trailing end of the rubjerg knude fyr section). this probably marked the end of the lateral translation along the lower flat levels and the initiation of stacking along steeply dipping thrust faults. this resulted in the rotation of all the previously formed structures and created the odd trough structures in the double ramp synclinal troughs. the final up-thrusting along steeply dipping thrust faults, which occurred contemporaneously with the uppermost sedimentation in the piggyback basins, was probably also contemporaneous with the initiation of ramping of st05 from the lowest level. one way to demonstrate this is to focus on the thrustfault development of st04. the st04 thrust fault acted as the ramp that pushed on the trailing end of the duplex below st01, st02 and st03. when the push on this footwall ramp ended and the st04 thrust sheet was displaced up over the footwall ramp of st03 it resulted in a shortening of 60%. the balanced length of st04 is c. 200 m. thus, the deep level st05 ramp must be responsible for removing the 120 m lower long segment originally situated below st04. therefore about half of the st04 thrust sheet also involves thrusting down to the 40 m décollement level. in the model, this st04u segment was up-thrust to form a first type of duplex-segment structure as the first lower-segment imbricate in the subsurface of st05. however, it should be appreciated that a whole unit of the thrust segments between the 20 m flat level and 30 m décollement level has to be incorporated in a differential thrust model. the simplest model for this is to dissect the lower duplex segments into sheets with an average length of c. 100 m. with each segment bend in a type 2 ramping and with an equal distribution of the frontal and trailing part on the upper and lower flat, a series of double ramp synclines would be created – comparable to the piggyback basins seen in st10, st09, st08 and st05. similar basins may have existed in st07 and st06, but, if present, were removed by glaciotectonic truncation. one of the central problems in describing the dynamic development of the rubjerg knude glaciotectonic complex is understanding the formation of the lower ramp below the st05 thrust sheet. it is known that the hanging-wall ramp is displaced up along the st05 thrust fault to be exposed in the cliff section at stortorn. however, a central question is – where was the footwall ramp for the lower hanging-wall ramp of st05 situated? according to the balanced cross-section, the st05 ramp should be situated about 7800 m from the frontal ramp in the ulstrup section and the footwall ramp for the lower st05 ramping should be situated on the far side (north) of the lønstrup village. however, this is not the position of the st05 ramp. the distance to a hidden footwall ramp can only be fixed relative to the displacement in front of the lower ramp when it was activated during the change of décollement level from the 30 m level down to the 40 m flat level. the relative displacement on the st05 thrust fault is c. 90 m, measured from the tip of the thrust sheet along the hanging-wall ramp-and-flat down to the lower décollement surface. the main problem is related to the compression documented south of the stortorn section. when this is considered in the balanced crosssection, it gives the geometric point for the hangingwall ramp 7800 m from the frontal ramp in the ulstrup section. however, the distance in the rubjerg knude cross-section from this ramp to the central part of the stortorn section is only 3800 m. the solution to this problem is that the ramp was first formed after all the former translation in the higher flat levels had passed. to understand this, one has to imagine that the upper part of the lønstrup klint formation at grønne rende (c. 3800 m from the frontal ramp in the ulstrup section) was originally situated above the stortorn formation at the st05 lower ramp. however, the ramp was first formed when the rubjerg knude fyr section was displaced towards the grønne rende section which itself was compressed against the stenstue and sandrende sections, which were all displaced over the initial position of the brede rende and kramrende sections. only then was the st05 lower ramp activated, and the remaining northern part of the rubjerg knude glaciotectonic complex was displaced along the lowermost décollement zone. 128 moserende section between rubjerg knude fyr and mårup kirke, some small peat-bogs occur in depressions in the dune landscape. this area of bog-filled depressions formerly extended to the west, and the present cliff section intersects one of these bogs where peat (martørv) is exposed in the uppermost part of the cliff, similar to the situation in the martørv bakker section. a former gully here was named moserende, and the name is adopted here for the section north of the stortorn section. the most impressive feature in the moserende section is the syntectonic evolution of the piggyback basins during polyphase thrust propagation (fig. 104). unusual sedimentological features are developed in the rubjerg knude formation, reflecting the tectonic deformation, notably structures described as fissure strata (sjørring 1977). these are thin sedimentary beds occurring as discordantly incised wedges in groundfrozen sediment, here present in the growth-fault synclines related to the piggyback basins (fig. 105). tectonic architecture the moserende section comprises 13 relatively thick thrust sheets annotated mr01–mr13. in the section, three larger piggyback basins are preserved, one in the frontal, southern part and two in the northern part of the section. the l/r-unconformity surface at the base of the piggyback basins was elevated to various levels in the cliff section, reflecting the differentiated type of ramping throughout the section. the leading-edge thrust in the moserende section is the hanging-wall ramp of mr01, which coincides with the footwall thrust fault on the back of the st10 thrust sheet. to the north, the section is bounded by the hanging-wall thrust at the base of the c. 40 m thick mk01 thrust sheet, which forms the southern front of the mårup kirke section. the mk01 thrust sheet was thrust up along a steeply dipping footwall ramp and subsequently displaced over the upper footwall flat on top of the piggyback basin of mr13 in the northernmost part of the moserende section. it should be noted that the main frontal part (topfig. 104. the piggyback basin of the mr02 thrust sheet. a sequentially developed growth-fault footwall syncline was formed below the hanging-wall ramp of mr03 (mr3hwr ). note the fissure strata cross-cutting the bedding (arrow) in the growth-fault footwall syncline. photograph: june 1984. 129 most part) of the thrust sheets shows a marked drag and truncation due to the formation of a glacitectonite on top of the cliff section. mr01 thrust sheet in the cliff section, the mr01 thrust sheet forms a massive unit of mobilised, structureless grey mud. the exposed thickness close to the beach level is nearly 30 m. the hanging-wall thrust dips at about 60°n, and the unconformity surface, which dips at about 45°, is elevated c. 15 m above sea level. this indicates that the mr01 thrust sheet is situated above two lower duplex segments. according to the balancing, these segments constitute a lower segment of the mr01 thrust sheet (mr01u) and a segment originating from the lowest, northern part of the stortorn section. the mr01u segment, which exists as a consequence of the estimated c. 46 m displacement along the 60° tilted frontal hanging-wall ramp, is separated into two differently displaced segments in the balanced cross-section. this is a feasible explanation but not the only one of several possible solutions for the displacement structure in the subsurface, which include differential lateral displacement along each lower 10 m level as well as mud diapirism. fig. 105. detail of the fissure strata indicated in fig. 104, illustrating that climbing ripple cross-laminated sands were deposited in the initially horizontal wedge-shaped fissure extending out into the growth-fault syncline deposits. photograph: june 1984; notebook (c. 18 cm long) for scale. 130 mr02 thrust sheet the mr02 thrust sheet is nearly 40 m thick and is dominated by mud mobilisation and internal chaotic structures reflecting polydiapirism and internal flow. the sheet is divided internally by a thrust-fault zone with dif ferential thrust movements, which could be interpreted as a separation of the sheet into two individual thrust sheets. however, as the segments are not separated by a piggyback basin they are regarded as a single amalgamated sheet. the displacement along the hanging-wall ramp is the same order of magnitude as for mr01 (c. 47 m) and the ramp is divided into an upper low-angle part with an initial dip of only 20° and a lower steeply dipping part with an initial ramp-angle of about 45°. due to subsequent rotation, the upper part of the ramp is now orientated vertically while the lower part has a steep listric dip to the south. the footwall ramp (upper part of mr01) has an initial dip of 45°, which creates a space problem in balancing the section and makes it necessary to introduce a splint segment between mr01 and mr02 in the subsurface. the splint was probably sheared and squeezed out and is likely to have been included in the general mud mobilisation. however, it is included in the balanced profile in order to deal with the ramp-angle-space problem (plate 2). the geometry of the l/r-unconformity at the top of the lønstrup klint formation in the mr02 thrust sheet is very irregular with a peculiar c. 8 m high obstacle. this is very similar to the structure in the st08 thrust sheet in the stortorn section. it was probably formed by a detachment anticline on the northern flank of the piggyback basin in the external part of the mr02 thrust sheet, where it was subsequently buried by the rubjerg knude formation sand. the unconformity is elevated about 5 m above sea level, indicating that the mr02 thrust sheet has only stepped up one level of the lower segments from where it is bent up along its hanging-wall ramp. the detachment anticline was probably formed during the ramping of this lower segment. mr03 thrust sheet the rubjerg knude formation in the piggyback basin on top of mr02 and mr03 is here envisaged as a single large basin situated in the frontal part of the moserende section. the basin extends about 70 m along the cliff section. three upright standing peaks represent the tips of three small thrust sheets that disturbed the basin by small displacements. these thrust sheets represent imbricates in the uppermost part of the mr03 thrust sheet. since the main part of the mr03 thrust sheet can be viewed as one large sheet subjected to a single mode of displacement, the three imbricates are referred to as mr03a, mr03b and mr03c. the l/r-unconformity in mr03b and mr03c can be traced down below sea level, indicating that the main part of mr03 was displaced along the lower décollement level prior to the displacement up along the 45° dipping frontal footwall ramp. however, the frontal hanging-wall ramp of mr03 is now vertical. it is only necessary to tilt the initial ramp on another 45° dipping ramp to achieve this, and although it is a very steep inclination for ramping, there are no obvious reasons for introducing more ramps. the steep ramp angle of the mr03b thrust forms part of the same framework. the initial dip of the mr03b hanging-wall ramp was 18°, and the thrust fault is now vertically orientated due to the ramp-bending mentioned above. the displacement relative to mr03a is only about 10 m and the sand beds of the rubjerg knude formation were folded in a footwall syncline of mr03a during the hanging-wall thrusting of mr03b. mr04 and mr05 thrust sheets the mr04 and mr05 thrust sheets are closely related and only separated from each other by a relative displacement of about 25 m along the hanging-wall thrust of mr05. in contrast, the displacement along the mr04 hanging-wall thrust is about 80 m. both thrust sheets are dominated by mud diapirism, structures that may have originated as one large diapir that was only displaced by the late mr05 hanging-wall ramp. the thickness of the lønstrup klint formation in the thrust sheets is up to 30 m in the cliff section, and the height of the diapir is 15 m. the diapir has characteristic intrusive contacts with the upper and frontal part of mr04 (fig. 106). it is evident that the final thrust displacement post-dates the diapirism, and the very steep thrust angle indicates that the ramping is rooted in the deepest levels of the section. the l/r-unconformity is elevated up to 10–12 m above sea level. from this it is inferred that the thrust sheets were lifted up on the lower segments of the mr03 and mr04 thrust sheets, although most of the lift is related to the ramping on the steep thrust faults. the thickness of the 131 rubjerg knude formation in the piggyback basin of the thrust sheets is about 15 m. mr06–mr08 thrust sheets the main feature of the mr06 and mr07 thrust sheets is that they are lifted relatively high up in the cliff section, such that only a small part of the piggyback basins are preserved. the ramping of mr06 and mr07 is about 22 and 13 m, respectively, corresponding to one level of elevation of the foremost mr06 thrust sheet. the ramping is interpreted to have been a stepwise progression up over the trailing lower part of mr05, which is ramp-bent over the lower segment of mr04. the displacement of mr06, mr07 and mr08 on each hanging-wall thrust fault is about 40 m, indicating that the displacement is of the order of the distance down to the décollement surface. the general impression is that mr06, mr07 and mr08 initially formed one large thrust sheet, which was stepwise separated during dif ferential thrust movements. this dif ferential thrusting moved mr06 to the highest position, whereas mr08 was left in the trailfig. 106. mud mobilisation and diapirism in the lønstrup klint formation in the central part of the mr04–mr05 thrust sheets of the moserende section. photograph: may 1985. fig. 107. mobilised mud (lower left) in the lower part of the lønstrup klint formation intruded into the bedding of the formation. the mobilised mud probably formed a viscous liquid that facilitated the gravity-spreading deformation mechanism. photograph: june 1984; staff divisions are 20 cm. 132 ing part with its hanging-wall flat still resting on the lower décollement surface. this is implied by the elevation of the l/r-unconformity, which can be traced down to a horizontal orientation about 5 m below sea level in mr08. in the frontal part of the mr08 thrust sheet, a very well-developed intrusive contact of a diapir is exposed (fig. 107). it is evident that the process of thrust faulting was facilitated by the buoyancy and lubricating effects of the water-saturated mud. in addition, the mobilised mud had the effect of pushing the thrust sheets from the rear during the gravity spreading process. mr09 thrust sheet the mr09 thrust sheet is a relatively thick thrust sheet. the angle between the main part of the hanging-wall ramp and the bedding in the lønstrup klint formation within the sheet is c. 30°, and the displacement is estimated to be about 58 m. the l/r-unconformity is elevated to about 7 m above sea level, and the mr09 thrust sheet must be considered to have been displaced along the rear part of the lower segments in the section. mr10 thrust sheet the mr10 thrust sheet has the same characteristic shape as the st08 and mr02 thrust sheets, with an obstacle interpreted as a detachment anticline. the final orientation of the hanging-wall ramp is rather steep (c. 70°) implying rotation of an initially steep ramp (c. 35°). this corresponds well with the angle between the bedding in the piggyback basin of mr09 and the mr10 hanging-wall ramp. in the cliff section, the lønstrup klint formation within the thrust sheet is up to 35 m thick. due to uncertainties in reconstruction of the frontal part, the displacement is estimated to be between 30 and 65 m. the l/r-unconformity can be traced down to about 5 m below sea level, indicating that the main part of mr10 rests on the lower décollement surface. the rubjerg knude formation of mr10 is about 25 m thick. it contains a c. 12 m thick lower unit with bedding dominated by r-onlap. above this follow three units, each developed as growth-fault footwall synclines. to obtain the rather large accumulated thickness of deposit in the piggyback basin, as well as folding the three synclines, the order of displacement is more likely to be 65 m than 30 m. this order of displacement also assumes that the frontal part of the thrust sheet extended nearly 50 m further ‘up in the air’ before being removed by erosion. the final implication is that the upper piggyback basin above the anticlinal obstacle was deposited in a syntectonically deeply eroded depression. mr11 thrust sheet mr11 is a small thin thrust sheet with a displacement of 55 m along the hanging-wall ramp. the thrusting of mr11 is another example of a thrust sheet requiring the formation of a splint in the subsurface. this is due to the low angle of the initial frontal ramp (only 15–18°), whereas the ramp angle between the hanging-wall ramp and the original bedding in the piggyback basin of mr10 is 40–45°. the splint was probably trapped as a triangular prism along the 45° dipping ramp, just below the beach surface. the l/r-unconformity is elevated to about 7 m above sea level. the model for balancing this thrust sheet indicates that it had a lower hanging-wall flat situated at the 30 m level (below the l/r-unconformity). from this level, it ramped up to the 20 m level onto a footwall flat on mr10, and finally developed an internal duplex of its own lower segment, which separated into mr11u 1 and mr11u 2 (see plate 2). the piggyback basin of mr11 consists of only a 10 m thick unit of the rubjerg knude formation. the sand beds show a f-bedding relationship to the l/runconformity with a weak tendency for d-onlap. mr12 thrust sheet the exposed part of the lønstrup klint formation in the mr12 thrust sheet is about 15 m thick and was displaced c. 58 m along a hanging-wall ramp dipping 38°. about half of the lønstrup klint formation here constitutes mobilised mud, without showing any marked tendency towards diapirism. the rubjerg knude formation is typically c. 15 m thick, with the exception of increased thicknesses in growth synclines, and the mr12 thrust sheet is regarded as only 30 m thick. this leaves another nearly 60 m long lower segment to be added to the trailing end of the mr10u duplex segment. the initiation of the mr12 thrusting started relatively early compared to the thrust sheets in front (to the south) and behind. this is based on a 133 consideration of the thickness of the rubjerg knude formation in mr11, which only reached a thickness of about 10 m before it became trapped by the mr12 thrust sheet. in the cliff section, the mr12 hangingwall flat (related to the 20 m flat level) is positioned on the footwall flat of the mr11 thrust sheet. the l/runconformity is situated about 5–6 m above sea level, which is compatible with the lift of the thrust sheet up onto the footwall flat of its own lower segment (mr12u). the rubjerg knude formation of mr12 was deposited in one of the large piggyback basins in the proximal part of the moserende section. the width of the basin is about 45 m and the accumulated thickness of the sand succession is c. 23 m. four units can be differentiated in the rubjerg knude formation of mr12. the lowest of these is about 10 m thick, displays fbedding, and corresponds well to the lower part of the rubjerg knude formation in other parts of the section. above this follows a trough-shaped unit folded in a gentle syncline. a new trough-shaped unit, which is folded into a tight footwall syncline, truncates the northern limb of this syncline. finally these two growth synclines are truncated by the upper sub-unit; the latter is mainly f-bedded, except for the northern part which is dragged into a footwall syncline below the hanging-wall thrust of mr13. mr13 thrust sheet the lønstrup klint formation of the mr13 thrust sheet forms an upright, wedge-shaped feature, which was displaced about 42 m up along a relatively steep (70° dip) hanging-wall ramp. this thrust surface has a remarkable curved shape, which is interpreted to be the result of erosion in the ramp caused by water flow contemporaneous with the deposition of the two growth synclines in the external part of mr12. similar features have been observed in places further to the south, but this is one of the best-developed examples. the erosion can be determined to have taken place in the interval between deposition of the lower 10 m thick unit of the rubjerg knude formation subsequent to c. 20 m ramping and before the final displacement along the ramp and deposition of the uppermost unit in the mr12 piggyback basin. the l/r-unconformity is situated 5–6 m above sea level. this is close to the elevation of the l/r-unconformity in mr12, and the thrusting follows a similar development style with ramping and displacement along the flat of the lower segment of the trailing end of the thrust sheet in front (to the south). the section balance requires a lower segment of the mr13 thrust sheet (mr13u) in the subsurface, bounded by the flats at the 20 m and 30 m levels (plate 2). the mr13u segment accumulated on the mr10u segment leads to a high ramping of the thrust sheet to the north, as will be demonstrated in the following section. the large piggyback basin in the moserende section is represented by the c. 20 m thick succession of the rubjerg knude formation in mr13. it is nearly 60 m wide and the main part is planar parallel bedded (f-bedded); only in the uppermost, rear part of the thrust sheet does a single footwall syncline appear. the external part of mr13 is therefore thought to have been deposited during a relatively long period of transport along a flat, contemporaneous with the ramping that had started in the thrust sheets in the distal part of the section. sedimentary units in the moserende section, only two formations involved in the thrust-fault deformation can be studied. the lønstrup klint formation forms the lower part of the thrust sheets, and the rubjerg knude formation forms the fill of the piggyback basins above the l/r-unconformity. lønstrup klint formation the lower part of the lønstrup klint formation typically consists of clayey mobilised mud. about twothirds of the thrust sheets comprise mobilised mud, which commonly developed into diapirs, rising from the sole thrust up into the formation, where they often truncate bedding with intrusive contacts (fig. 107). the upper part of the formation is composed of thinto medium-bedded sandy turbidites interlayered with silty mud (fig. 27). in the majority of layers, the mud and fine-grained sands have been disturbed by balland-pillow breccias or small polydiapiric water-escape structures. during the main thrust faulting, these structures were superimposed by a dense framework of smaller thrust faults (fig. 27). 134 rubjerg knude formation the rubjerg knude formation has a maximum thickness of 25 m in the moserende section; such thicknesses are rarely attained, however, either due to overthrusting that sealed the deposits in the piggyback basins before accumulation of this thickness, or to erosion of the upper part of the formation after deposition and deformation. in general, deposition was initiated with a unit of f-bedded sand c. 10 m thick. above this are two or three units showing r-onlap and locally foreland-dipping large-scale cross-bedding (d-onlap). the uppermost part typically shows r-onlap. structures in the moserende section, three types of structural elements are described and discussed: (1) diapirs, including mud mobilisation, (2) mesoscale thrust faulting, and (3) growth-fault footwall synclines, including fissure strata. the existence of frozen sand clasts and frost wedges that testify to the ground-frozen condition of some of the sediments in the rubjerg knude glaciotectonic complex has already been mentioned. further evidence for ground-frozen conditions is seen in the presence of fissure strata (sjørring 1977). a fissure stratum is a layer of sand deposited horizontally in a fissure that discordantly cuts into a package of sediment (usually meltwater sand) (fig. 105). the implication of the occurrence of fissure strata is that not only was the host sediment affected by (glacio)tectonic deformation prior to fissure incision, but also that the sediment must have been frozen so that the fissure cavity did not collapse during deposition of the fissure strata. as the fissure strata form wedge-shaped sand layers, they have also been referred to as kilelag (wedge-layers in danish; berthelsen 1975). in the moserende section, the fissure strata document an intermediate phase of syntectonic deposition, as they have been tilted into a vertical position. diapir structures in the moserende section, the zone above the hanging-wall ramps and flats often comprises mobilised mud (figs 106, 107). the mud-mobilisation and related diapirs dominate in the thrust sheets from the intermediate hanging-wall ramp and towards the trailing end of the sheets. the diapirs are irregularly developed, mainly related to intrusive migration laterally into the bedding. in the example shown in fig. 107, the intrusive mobilised mud has a contact rim of segregated sandy mud that forms the contact to the truncated bedding. in the thickest thrust sheets, the larger diapirs rose from the hanging-wall thrust fault up to 15 m above the thrust-fault surface (fig. 106). in the construction and calculations of the balanced cross-section (plate 2), the mobilised mud and polydiapirism create a problem of volume preservation relevant to the approximation and evaluation of the reliability of the balanced model. however, the volumes are not lost but only reorganised and may therefore be treated as part of the thrust sheets and duplex segments. the mobilisation is dominantly developed along the lower part of the thrust sheet, from where the mud intruded the upper part of the lønstrup klint formation of the thrust sheets. it is evident that the thrust sheets were carried on the mobilised mud, which with its high water pressure facilitated the displacement of the sheets. pedersen (1987) described a model for this process, and with minor modifications this is still considered to be valid (fig. 4). the model also implies that the muddy liquid formed a pressure agent in the gravity-spreading dynamics, which pushed the thrust sheets forward towards the distal part of the thin-skinned thrust-fault complex. thrust faults in the structural cross-section, only the thrust faults identified as carrying the major displacements are outlined (plate 1). a number of smaller thrust faults and bedding-parallel contractional faults that occur in the rubjerg knude glaciotectonic complex are therefore not included in the cross-section. one example of these less significant faults was described in the rubjerg knude fyr section (fig. 101). in the moserende section, similar thrusts occur, and were recorded during detailed logging of the upper part of the lønstrup klint formation in the mr03 thrust sheet (fig. 27). within a 7 m thick succession, ten minor thrust faults have been recognised. the base of the succession is the hanging-wall ramp of mr03b. along the base of the succession, 0–1.5 m above the hanging-wall ramp, the clayey mud is cataclastically brecciated by anastomosing shear fractures. in the overlying part of the formation, the bedding-parallel thrust faults occur with a spacing of 0.5–1.5 m, concentrated in the muddy 135 layers, whereas steep connecting ramps are situated in the sandy beds. it is likely that this type of differential beddingparallel thrust faulting and shear brecciation was an important component in translation in the thin-skinned thrust-fault system. it also implies that the stratigraphic succession in a formation that appears to be well preserved may in fact have experienced significant lateral dislocation. footwall synclines the dominant structures in the moserende section are the footwall synclines, which include the growth syncline basins and re-orientated fissure strata. these synsedimentary folds were formed continuously, beginning with the deposition of the sand in a trough. as the thrusting up over the footwall ramp progressed, the trough deepened; contemporaneous with this deepening of the footwall block, the hanging-wall ramp north of the trough started thrusting, which resulted in a drag bend of the northern flank of the syncline. some of the synclines were trapped and overthrust, resulting in overturning of the northern flanks, which in a few cases created nearly isoclinal, recumbent folds. in the piggyback basin of the mr02 thrust sheet, an excellent example of deposition in a growth syncline is preserved. the width of this basin (c. 20 m) is of the same magnitude as the thickness of the sequence deposited and deformed. the basin fill can be divided into four sub-units of the rubjerg knude formation separated by angular discordances (fig. 104). the lowest sub-unit (s 1 ) consists of c. 5 m of trough crossbedded sand deposited in an erosional depression incised into the lønstrup klint formation of the mr02 thrust sheet. it could be argued that in relation to the l/r-unconformity of mr02, this sub-unit shows r-onlap, but it is evident that the trough cross-bedded sand is growth-related to an initial up-thrusting of the mr03 frontal nose. the second sub-unit (s 2 ) was deposited after the first sub-unit was tilted during ramping of the mr02 thrust sheet. the tilting also lifted the s 1 sub-unit up to a level at which the 5 m thick sand package could be subjected to ground-frost. this is deduced from the occurrence of fissure strata emplaced discordantly into the s 1 sub-unit. the fissure strata transecting sub-unit s 1 show small-scale current ripples, which demonstrates that these fissure strata were deposited parallel to the sides of the wedge (fig. 105). the fissure strata can be traced into the second sub-unit s2, which consists of stratified sands characterised by climbing ripple cross-lamination, deposited in a growth syncline trough. the thickness of the s2 sub-unit is 5– 10 m; this sub-unit is dominated by r-onlap in relation to the l/r-unconformity in mr02. deposition of the third sub-unit (s3) was first initiated after sub-units s1 and s2 were deformed by compressional deformation due to push from hangingwall thrusting of mr03 over the footwall block of mr02. the fold structure may be described as an inclined s, where the lower bend of the s corresponds to a footwall syncline. the s3 sub-unit is c. 10 m thick. the base of s3 is an angular discordance on the folded s1 and s2 sub-units. the s3 sub-unit shows r-onlap onto the l/r-unconformity of mr02 and the uppermost part of this sand package covers the detachment anticline structure of mr02 as well as filling the depression on the northern flank of the structure. finally the s3 sub-unit was folded into a footwall syncline by thrust propagation of mr03. in the uppermost part of the basin, a small growth syncline comprises the uppermost sub-unit (s4). this unit is 1–6 m thick and may be regarded as recording deposition in a depression formed by a footwall synclinal bend during the general northward tilting of mr02. interpretation of structural development in the model for the structural development of the moserende section, a distal (southern), an intermediate and a proximal (northern) zone are distinguished. the distal zone includes thrust sheets mr01–mr03, which were probably displaced in similar mode. the intermediate zone includes thrust sheets mr04–mr08, and finally in the proximal zone thrust, sheets mr09– mr13 probably moved sequentially in one continuous displacement. the distal mr01–mr03 zone is regarded initially to have formed one coherent thrust sheet, which was split up during the ramping of mr01. the main bend during ramping took place in mr02, while the trailing end of mr03 still rested on the décollement surface without being elevated, and thus represents the root zone of the mr01–mr03 segment. the sequential development of deposition and deformation is very well illustrated in this first segment with the key locality in the piggyback basin of mr02. here four depositional 136 phases (s1–s4) separated by four deformational phases (f1–f4) have been identified. depositional phase s1. the s1 depositional phase took place during the initial thrusting. about 5 m of trough cross-bedded sand was deposited before (or coeval with) the first deformation phase f1. deformational phase f1 . the first fold phase culminated with the creation of the hanging-wall anticline or detachment anticline in mr02. this folding must be the effect of ramping up from the 20 m level (or 30 m) to the 10 m flat level, resulting in folding of the c. 10 m upper lønstrup klint formation of mr02. depositional phase s2 . during s2 deposition, the angle between bedding in the rubjerg knude formation and the l/r-unconformity is nearly perpendicular. the ramp is not indicated in the ramp cross-section because it was destroyed by mud-mobilisation in the lower part of mr02 (plate 2). deposition of s2 also took place during the exposure of the s1 unit to ground frost conditions as indicated by the presence of fissure strata. deformational phase f2. the f2 folding was related to the hanging-wall thrusting of mr03. as this phase includes thrust displacement of the s1 sand and a synclinal drag of the s 2 sand below the footwall ramp of mr02, it progressed during s 2 deposition. the f 2 phase terminated with the final folding and minor thrust truncation of s 2 . thus the separation of the imbricate thrust sheet mr02 and mr03 took place during the f 2 phase. depositional phase s 3 . deposition of s 3 covered the hanging-wall anticline of mr02, and part of s 2 is discordantly inclined relative to s 3 bedding. s 3 sedimentation is characterised by f-bedding, and displacement probably took place along the 30 m level on top of mr01u, and also st09–st10u. deformational phase f 3 . the f 3 fold phase is restricted to folding of a footwall syncline of the s 3 unit related to the hanging-wall ramp propagation of mr03. it might be interpreted as a continuation of f2, but the mr03 thrusting definitely propagated after s 3 deposition and its extension most likely forms a hangingwall flat over the sand covering the hanging-wall anticline of mr02. depositional phase s 4 . the uppermost growth syncline in the piggyback basin of mr02 was formed by the 5 m thick uppermost unit of the rubjerg knude formation sand. it probably truncated the f3 thrust, and is thus not interpreted to represent dislocated parts of s 1 or s 2 . deformational phase f 4 . the last fold phase created the footwall syncline of the s4 deposits. the hanging-wall ramp of mr03 increased its inclination by about 30° and propagated along a steeper satellite thrust. the footwall syncline was initially overturned to the south, but during the final phase of ramping and steep tilting the syncline became re-orientated into an upright position. the mr04–mr08 intermediate zone initially formed one coherent thrust sheet, comparable to the mr01– mr03 thrust sheets, with a frontal steep ramp, hanging-wall ramp of mr04, and a mr08 trailing-end sheet with the l/r-unconformity below sea level, indicating that this thrust sheet rooted in the décollement zone. in the proximal zone, the mr09 thrust sheet probably propagated as an individual sheet onto the footwall ramps. the main thrusting was initiated with the mr11–mr13 sheets being thrust along a flat in the 30 m level, which is on the lower footwall flat of the trailing end of the mr10 thrust sheet. thus mr11– mr13 can be viewed as thrust-sheet imbrications peeling off the back of mr10. at the latest stage of thrusting, the continued displacement along the décollement zone was responsible for the steepening up of the sheets, somewhat similar to the development of the grønne rende section. for the calculation of the displacement of the individual thrust sheets, the erosionally removed frontal parts have been reconstructed from simple angular geometry, in the same way as the calculation of displacements in the grønne rende section (fig. 11). when the elevation of the l/r-unconformity is considered, it mainly refers to the position of the lowest part of the l/r surface. ideally this part should be horizontal, to indicate the main elevation of the thrust sheet. however, this is not always the situation, and therefore the position of the l/r-unconformity in general refers to its position where a hanging-wall ramp or flat truncates it. mårup kirke section one of the locations where landsliding at present is most dramatic is at mårup kirke (the old church at mårup). this attracts much public attention, not least among the local people. the back-stepping of the head of the slides has been very rapid during the last ten years, and it is probable that the coastal protection measures instituted at lønstrup have increased the erosion of the clif f at mårup kirke. over a number of 137 years the landslides have also obscured the exposures at this location and a lot of interpretation has been necessary to reconstruct the structural framework and the development of the thrusting. the younger yoldia clay and saxicava sand (the vendsyssel formation) cap the mårup kirke section. the erosional unconformity at the base of the vendsyssel formation acts as a drainage surface, which contributes to the generally poor exposure conditions. however, with the experience gained from the other sections, combined with theoretical structural analysis, it is possible to present a model of the structures in the mårup kirke section. there is a marked correlation between the occurrences of piggyback basins comprising the sand-rich rubjerg knude formation, and the build-up of aeolian dunes above the cliff. the northern boundary of the dune field is situated about 200 m south of mårup kirke in the frontal part of the mårup kirke section. south of this area, the aeolian dunes form features ranging from a few metres in height to nearly 50 m high dunes above the rubjerg fyr section, which is also the highest point of the cliff section. thus it is evident that the dunes are formed where a sand source (the rubjerg knude formation) is available (pedersen 1986b). above the main part of the mårup kirke section, no dunes are present, as this is a cliff section that consists only of clay and mud (fig. 108). the most interesting feature at mårup kirke is the packing of the thrust sheets into uniformly developed thrust-fault deformed duplex segments. a theoretical model is presented for the deformation geometry of these duplex units, subjected to extreme compressional development; comparisons with the observations refig. 108. the mårup kirke situated at the head of the cliff in the central part of the mårup kirke section. note that the flat cliff-top surface, representing the horizontal bedding of the vendsyssel formation, is not covered by dunes, reflecting the fact that the clif f consists of clay and mud. in the distance, the rubjerg knude fyr (lighthouse) is being engulfed by sand dunes derived from the sand-rich rubjerg knude formation in the piggyback basins present in the sections to the south. photograph: july 1994; note that by 2002 cliff erosion had reached the corner of the graveyard. 138 corded in the clif f section indicate a reasonable match between the theoretical model and cliff observations. tectonic architecture due to the variation in the distribution of piggyback basins and the general tectonic architecture of compressional framework, the mårup kirke section is divided into three zones: (1) a leading zone (mk01– mk07), (2) a transitional or intermediate zone (mk08– mk10), and (3) a trailing zone, including thrust-fault duplex units annotated mk11–mk20. in the leading zone, the first six thrust sheets have preserved a relatively highly elevated remnant of their strongly eroded piggyback basins. in the transitional zone, only the thrust sheet mk10, just below the mårup kirke, contains a piggyback basin. north of this thrust sheet, the thrust-fault duplex units only comprise the lønstrup klint and stortorn formations. in the intermediate and trailing zones (mk08–mk20), the thrust sheets are defined as duplex units, each comprising four duplex segments, which are annotated d1–d4. thus the lower duplex segment in unit mk09 is annotated mk09d1 and the upper thrust-sheet segment is annotated mk09d4 (plate 2). the southern boundary of the mårup kirke section is the leading-edge thrust fault corresponding to the trailing footwall thrust of mr13 and the hanging-wall ramp of mk01. the boundary is fairly obvious, as it separates the large piggyback basin of mr13 from the c. 50 m thick mk01 thrust sheet. the trailing end of the section is more loosely defined, due to poor exposure and the increasing mud mobilisation of the thrust sheets. it has therefore been defined in relation to the unconformities above the thrust sheets. thus the boundary between the mårup kirke section and the ribjerg section to the north is placed where the fig. 109. upright thrust sheets in the frontal zone of the mårup kirke section (thrust fault arrowed). the dif ference in lithology between the upper and lower levels of the lønstrup klint formation is illustrated by comparing the sand-rich (upper lønstrup klint formation) footwall block (to the right, south) with the mud-dominated (lower lønstrup klint formation) hanging-wall block (left). encircled rucksack for scale. photograph: august 1996. 139 unconformity at the base of the vendsyssel formation truncates the unconformity between the glaciotectonic unit and the ribjerg formation. the glaciotectonic unit is viewed as the unit of deformed deposits related to a glaciotectonic event (pedersen 1993), here included in the rubjerg knude glaciotectonic complex. the defined boundary is very close to where 45° dipping thrust-fault features are obscured both by the mud mobilisation and by superimposed, more or less horizontal, anastomosing jointing. mk01 thrust sheet the first thrust sheet in the leading zone of the mårup kirke section is the nearly 50 m thick mk01 thrust sheet. in the exposed parts of the thrust sheet, the lønstrup klint formation has a thickness of 40 m, which indicates that the deepest level of the section including the lower décollement zone is brought up to the upper flat. the displacement is estimated to be 102 m; this includes construction of the eroded tip of the thrust sheet (fig. 11). this corresponds well with calculation of the displacement according to the equation d × sinα = h, where d is the displacement, α is the angle of the thrust ramping from the lower flat to the upper flat, and h is the distance between the two flats. the distance between the lower and upper flats is 40 m, and assuming an initial thrusting angle of about 23°, the calculated displacement d is c. 100 m. most of the lønstrup klint formation of mk01 is mobilised and forms a diapir-like structure. bedding is only well preserved in the upper 7–10 m of the lønstrup klint formation in this thrust sheet. the bedding is characterised here by medium-bedded sandy turbidites. the rubjerg knude formation of mk01 comprises a nearly 10 m thick succession of glaciofluvial sand. the sand layers show r-onlap and are strongly deformed by a footwall syncline. the l/r-unconformity is situated about 15 m above sea level, indicating an elevation of the lower hanging-wall flat up to the 20 m level footwall flat. this corresponds well with the balanced model for the moserende section, where the trailing-end lower segments mr10u and mr13u still remain to be calculated for. it is thus evident that the mk01 thrust sheet was translated along the 20 m level flat. mk02–mk04 thrust sheets the mk02–mk04 thrust sheets are three relatively small sheets with only minor displacements, about 30 m for mk02 and mk03, and c. 40 m for mk04. the estimates of the displacement for thrust sheets with erosionally removed tips are based on reconstructions following the same principles as applied in the grønne rende section (fig. 11). the thickness of the lønstrup klint formation in the thrust sheets is only between 10 and 20 m, and the thrust sheets are considered to represent imbricates from the upper part of the mk01 thrust sheet, which roots down to the décollement zone in the 40 m flat level. the lønstrup klint formation in mk02 is completely mobilised apart from the uppermost 1–2 m. the mud diapir in mk02 is regarded as an extension of the large mud diapir in mk01. in mk03 and mk04, bedding is partly preserved, and may be compared with the upper sandy bedding seen in mk01. the l/r-unconformity cuts down to at least 2–3 m below the mean level of bedding on the back of mk02 and mk03, and given this uncertainty in location, the position of the unconformity is at the same elevation as in the mk01 and mk04 thrust sheets. it is therefore inferred that the imbricate thrusting of mk02–mk04 was initiated at an early stage, prior to the subsequent deeper-level up-thrusting of mk01. early development of the imbrication is further supported by the thickness of the rubjerg knude formation, which in mk01–mk04 is less than 10 m. the piggyback basin on the back of the non-imbricated mk01 was short-lived relative to those described above in the moserende section where the rubjerg knude formation is up to 25 m thick. along the footwall ramps, sand of the rubjerg knude formation was deposited in growth synclines indicating syntectonic development of the small piggyback basins in the leading zone of the mårup kirke section. mk05–mk07 thrust sheets the mk05–mk07 thrust sheets form an imbricate set of upright thrust sheets, now dipping more than 60°n (fig. 109). the thickness of the thrust sheets is about 20 m in the cliff section. the displacement is 40 m for mk05, 28 m for mk06 and only 15 m for mk07. the lønstrup klint formation, making up most of the thrust sheets, is characterised by a few relatively thick finegrained sand turbidites interbedded with dark blue140 grey silty mud. the bedding is strongly disturbed by thrusting and jointing, similar to the thrust-fault framework described in mr03, and smallto medium-scale duplexes are common. in the upper half of the mk07 thrust sheet, a major footwall syncline is developed below the hanging-wall ramp of mk08. the rubjerg knude formation is only represented in mk06, where the l/r-unconformity is elevated up to about 20 m above sea level. in the mk04–mk05 thrust sheets the elevations are more than 20 m, indicating that these thrust sheets were lifted up and translated along an upper flat resting on three lower segments. the accumulation of these segments forms a subsurface duplex, probably deformed into an intense network of anastomosing thrust faults. the duplex segments constitute the trailing lower segments of mk01–mk04, and for mk06 and mk07 the trailing lower segment of mk05 is also added (see balanced crosssection, plate 2). mk08–mk10 thrust sheets mk08 is the southernmost thrust sheet in the transitional zone of the mårup kirke section. the transitional zone is characterised by the change from thick continuous successions of lithologies into sheet units comprising duplex segments bounded by footwall and hanging-wall flats. along these flats, lateral translation preceded tilting and steepening during propagation along the ramps. the thrust fault separating mk07 and mk08 is the trailing-end footwall ramp on top of mk07 and the hanging-wall flat of mk08. at the top, mk08 is bounded by the footwall flat and hanging-wall ramp between mk08 and mk09. the thrust fault and the general bedding in mk08 dips at 35°n. in the exposed cliff section, mk08 is 40 m thick, which implies that the hanging-wall flat is a segment of the lower décollement zone. above this, four segments may be identified corresponding to the four main levels of flats below the l/r-unconformity. each flat-segment is about 10 m. the displacement along the hanging-wall thrust is c. 110 m, calculated from the equation d × sinα = h, given that the height h is c. 65 m (sum of cliff section and distance down to the décollement surface) and α is 35°. the estimate is based on the assumption that the tip of the hanging-wall ramp in the lowest segment (initially located at the 30 m flat level) only reaches up to the hinge of the footwall ramp at the top of the cliff. in the mk08 thrust sheet, there is no record of the rubjerg knude formation. due to the intense development of landslides below the mårup kirke, the mk09 thrust sheet is very poorly exposed. the interpretation here is based on the space relationships and the structures exposed in mk08 and mk10. from mk08 it is known that the frontal hanging-wall ramp of mk09 dips at 35°. the bedding is more or less horizontal in the cliff section, judging from the occasional features that can be picked out from the photo-geological interpretation. the distance between the leading and trailing thrust fault is about 70–73 m, and the displacement must therefore be about 50–55 m. the interpretation of the mk09 framework is that the upper segment forms a type 3 fault-bend-fold structure, where the sub-segment resting on the upper flat has been eroded away and truncated by the mk10 hanging-wall flat, the two segments in the intermediate levels form s-shaped type 2 structures, and the lower segment forms a type 1 structure with the trailing end of the segment resting on the lower flat (compare with the model in fig. 103). the mk10 thrust sheet is situated below the mårup kirke, and preserves the most proximal piggyback basin containing the rubjerg knude formation. the thickness of the rubjerg knude formation is about 10 m and the basin is deformed into a recumbent footwall syncline. the l/r-unconformity is situated about 10 m a.s.l., indicating elevation above two lower duplex segments in the subsurface. the increase in the dip of the l/r-unconformity and bedding in the lønstrup klint formation from 20° in the frontal part to 35° in the rear part is interpreted as a bend by the hanging-wall ramp propagating over irregularities in the trailing part of mk09. thus part of the subsurface structure must include the geometric adjustments of a splint appearing due to the ramp angle change (see plate 2). this is reflected in the occurrence of a hanging-wall anticline in the trailing end of mk10. mk11–mk20 thrust sheets in the trailing zone of the mårup kirke section, there are no occurrences of the rubjerg knude formation, and the l/r-unconformity has not been identified. the thrust faults are mainly steeply dipping, about 45°, and at the top of the cliff section the thrust sheets are shear-dragged southwards and reworked into glacitectonites, a truncated glaciotectonic unconformity and local till. 141 this trailing zone can be divided into ten fairly uniform thrust-fault duplex units, each about 55 m long (measured horizontally along the beach level) and separated by 45° steeply dipping thrust faults. for the characterisation and structural explanation of these thrust-fault duplex units, a thrust-ramp-propagation model is presented below (see fig. 114). sedimentary units in the mårup kirke section, no significant additional data have been obtained to supplement the sedimentological descriptions. however, it is evident that the rubjerg knude formation only occurs in the southern part of the section, where it is less than 10 m thick. it is inferred, therefore, that the piggyback basin in the mårup kirke section was short-lived relative to the thicker successions of the rubjerg knude formation further south. the rubjerg knude formation may never have been deposited in the trailing end of the mårup kirke section. structures a conspicuous feature of the mårup kirke section is the shear drag at the top of the cliff section. the main structures formed during the subglacial drag are southerly overturned to recumbent synclines that developed in a sandy glacitectonite about 1 m in thickness. the glacitectonite is interpreted to have formed by subglacial shear deformation superimposed on the proglacially formed thrust-fault and duplex structures (pedersen 1988, 1996, 2000). over a large part of the mårup kirke section, the amount of mobilisation is not very high. this permits a characterisation and interpretation of the deformation of duplex segments and stacking of duplex units, as described below. interpretation of structural development the main purpose of this section is to present an analytical structural model that can be used to interpret the thrust-fault framework developed in the mårup kirke section. the basic elements of the model are the duplex segments, and the structural deformation can be characterised as fault-bend folding. the result of the deformation is a compressional stacking of duplex segments into duplex units with a certain geometry and size. the interpretation of these structures adds to the basis for the discussion of structural developments that concludes this section. fault-bend-fold model for duplex units it was demonstrated in fig. 103 how a duplex unit developed with type 1–3 duplex-segment structures. however, a number of structural configurations may develop from the deformation of duplex segments stacked into duplex units, depending on the amount of displacement and the initial length of the thrust sheet. for the interpretation of the structures in the northern part of the mårup kirke section, as well as a major part of the ribjerg section, the analytical models in fig. 110 have been constructed. the premises for the models are: (1) the duplex unit comprises four initially horizontal sheets with a thickness of 10 m each, (2) the bounding leading and trailing thrust ramps dip at 45° (maximum angle of thrust-fracture formation), (3) the vertical distance between the lower and upper flat is 40 m, and (4) the lateral compression cannot exceed the packing of the sheets in 45° dipping imbricates. from the latter premise, it can be predicted by simple trigonometric calculation that the lateral distance between the bounding thrusts of the duplexes should be close to 56.5 m, and this corresponds very well with the thrust features recorded in the cliff section. to illustrate the model, one ideal case is considered, namely the case where the displacement is to the top of the ramp, which has the same length as the maximum compression distance of 56.5 m (fig. 110, type 3). the displacement takes place along the lower footwall flat, and the lower hanging-wall flat propagates up along the 45° dipping footwall ramp. the resulting structural framework is an l-type fault-bend folding (fig. 103). in this case, the lower thrust segment will just reach the level of the upper flat. the displacement is c. 42 m and the balanced length of the duplex unit is 98.5 m, which results in a calculated compression of 43%. due to the propagation along the upper flat and the ramp-bend folding, a hangingwall anticline is formed, which can be described as a fairly upright, angular antiformal stack. the case described above is shown as type 3 in fig. 110. this case might also be called the angular antiformal stack type. further cases can be considered with decreasing or increasing displacement relative 142 l0 (m) 70.5 1 20%compression 2 33% 3 43% 4 50% 5 56% 6 60% 84.5 98.5 113 126 140 7 l1 56.5 comp. = 28 84.5 33% l0δ 1 l1 56.5 comp. = 14 70,5 20% l0δ 2 l1 56.5 comp. = 28 84,5 33% l0δ 3 l1 56.5 comp. = 42 98.5 43% l0δ 4 l1 56.5 comp. = 56.5 113 50% l0δ 5 l1 56.5 comp. = 70.5 127 56% l0δ 6 l1 56.5 comp. = 84 140.5 60% l0δ 8 l1 56.5 comp. = 56.5 113 50% l0δ 9 l1 56.5 comp. = 70 126.5 56% l0δ fig. 110. the duplex-unit model for fault-bend folding of duplex segments. the basic elements for the constructed models are: (1) the duplex unit comprises four initially horizontal sheets, separated by thrust-fault flats, and each sheet is 10 m thick, (2) the bounding leading and trailing thrust ramps dip at 45° (maximum angle of thrust-fracture formation), (3) the vertical distance between the lower and upper flat is 40 m in types 1–6, and in types 7–9 it is extended 10 and 20 m above the upper footwall hinge, and (4) the compression cannot exceed the packing of the sheets in 45° dipping imbricates. the cases in the model are selected with steps jumping one 10 m level from case to case. according to simple trigonometry, this will result in a displacement unit of c. 14 m, and multiples of this. the initial dimensions (l 0 ) of duplex units 1–6 are given by the scale in the lower right corner. type 1 is a single monoclinic flexural kink fold. type 2 is a double monoclinic flexural kink fold. type 3 is an angular antiformal stack. type 4 is a flat-topped antiformal stack. type 5 is a lateral extension of the flat-topped antiformal stack. type 6 is a lateral extension of the flat-topped antiformal stack, where it is demonstrated that the frontal limb in the antiformal stack retains its profile, and it is only a lateral translation of duplex segments that responds to the further compression of the duplex unit. type 7 is a per fect g-s-l structure (fig. 107). compression of 33% in type 7, and an elevation of the ramp by two 10 m levels, results in the monoclinic flexural kink fold. the ef fect of increasing compression and propagation up along the extended ramp (types 8 and 9), demonstrates the development of the antiformal stack in a manner comparable to that from type 3 to type 4. note that the given maximum stacking of imbricates constrains the size of the balanced length of duplex units. this is demonstrated in the diagram relating balanced length to magnitude of compression. l1, length after deformation; δ, shortening; lo, initial length; comp., compression. to type 3. if the displacement of the lower thrust-sheet segment is less than the height of the footwall ramp, the duplex segments above have two ramps to pass. consequently, two ramp-bend folds will be created, which may also be described as a repeated monoclinic flexural kink-fold (fig. 110, type 2). in the model, the second ramp-bend fold will not develop until the displacement exceeds 20%, corresponding to a displacement lift of only one 10 m level (fig. 110, type 1). with increased compression, the hanging-wall anticline formed in type 3 will develop into a flat-topped antiformal stack (types 4 and 5). finally, type 6 dem143 onstrates the lateral extension of the flat-topped antiformal stack resulting from 60% compression. note that in this case the frontal limb in the antiformal stack maintains its profile and an increase in compression only results in lateral translation of thrust sheets. increasing compression also requires increasing length (l0) of the duplex unit, which is demonstrated by the diagram in the lower right corner of fig. 110. in fig. 110, the cases with an extended ramp have also been examined. this corresponds to thrusting above the upper hinge of the footwall ramp, which would be the case if syntectonic sedimentary units were deposited on the upper flat preceding ramp propagation. type 7 is a perfect γ -s-l structure, exemplifying this development. it is formed by compression of 33% and elevation of the ramp by two 10 m levels, here creating a monoclinic flexural kink-folding. the effect of increasing compression and elevation of the ramp from type 8 to type 9 demonstrates the development of the antiformal stack in a manner rather similar to that from type 3 to type 4. characterisation of thrust duplex mk11–mk20 on the basis of the models in fig. 110, the mk11 thrust sheet is classified as a type 2 structure due to the presence of two monoclinal flexures. however, the structure in mk11 must incorporate the effects of the displacement of the thrust segment of mk10. consequently, the upper segments of mk11 are stacked on each other as relatively short duplex segments. moreover, the topmost part of mk11 is dragged out and sheared over the piggyback basin of mk10. this dragged part can be interpreted as the frontal limb of the antiformal stack initially formed over the upper footwall hinge. mk12 is the duplex unit situated north of mårup kirke. all the structures dip at 45°, except for the uppermost shear-dragged parts, which were reworked into a local till (the kattegat till formation). thus the structure is interpreted mainly as an l-structure, probably a type 5 or 8 structure with 55% compression and a balanced length of c. 126 m (fig. 110). mk13 has an undulating flat-lying structure with flexural drag up along the footwall ramp. it is thus interpreted as a type 4 structure with a flat-topped antiformal stack capping the frontal part of the lower thrust segment, which was only displaced up to the reference level of the l/r-unconformity. compression amounts to 50–55%, and the balanced length is estimated to be 120 m. mk14 is considered to be similar to mk13. it was probably very close to the modelled type 4 structure (fig. 110), prior to glaciotectonic shearing and truncation of its flat-topped antiformal stack. mk15 and mk16 are probably the closest approximation to a perfect γ-s-l-structure of type 7 in the model (fig. 110). mk17 and mk18 may well be inferred to be of the same type. however, the exposures are here too poor for definitive structural characterisation. in mk19 and mk20, structures with 45° steep dips are displayed in the cliff section. these duplex units can thus be interpreted as type 9 duplexes. discussion of structural development although the balanced section is subject to some uncertainties in the mårup kirke section, calculation of the compression from the measured length of the section l1 = 978 m, and a balanced length of about l0 = 1814 m gives 46%. as described above, the section is divided into three architectural zones: (1) a leading zone (mk1–mk7), (2) a transitional or intermediate zone (mk8–mk10), and (3) a trailing zone that includes thrust-fault duplex units (mk11–mk20). the discussion below attempts to demonstrate the proximal–distal thrust-fault development. the fault-bend-fold model for duplex units describes the thrust-fault structures in the trailing zone and gives an approximation of the structural framework of the major part of the mårup kirke section. the absence of the rubjerg knude formation in the trailing zone suggests that it was never deposited here. moreover, the thrust stacking of the duplex units started before, or just at the beginning of, deposition of the rubjerg knude formation in the most proximal part of the glaciotectonic complex. it is further suggested that the thrust levels rapidly shifted to lower levels in progressive steps. so, after the first few hundred metres of peeling off the uppermost thrust segments, the thrusting propagated for the next five hundred metres in the intermediate flat levels. finally, the main compression started to stack the duplex units up into imbricates during translation along the lower flat level, the décollement zone, and differential displacement between the duplex segments. the displacements of the duplex units were limited by the maximum shortening between the 45° steep northward-dipping ramps. it might be suggested that the displacement was much larger and considerable amounts of the leading part of the thrust sheets were 144 eroded away from the upper flat. however, this is unlikely for two reasons: (1) the amount of compression is 50–60% which is considered to be a limiting amount of compression for natural systems, and (2) the structures discernible from the photo-geologically interpreted cross-section support a model with c. 50% shortening. another suggestion could be that the duplex imbricates were formed subglacially, bounded by a floor thrust (the décollement zone) and a roof thrust situated in the glaciotectonic unconformity (the sole of the glacier). this is disproved by the fact that the antiformal stack above the duplex units would have required space to be stacked up on the upper flat. thus, although the antiformal stacks were removed by glacial erosion and the upper part of the mårup kirke section is shear-dragged and truncated by the glaciotectonic unconformity (formed subglacially), the thin-skinned thrust faulting developed in a proglacial setting in front of a progressively advancing ice margin. mk08 is the leading thrust sheet in the transitional zone of the mårup kirke section. it has a considerable displacement, more than 100 m, and it probably ramped up to the 20 m flat level along which it was translated for more than 50 m before its hanging-wall flat propagated up to the uppermost footwall flat. thus, all the segments in the thrust sheet were earlier translated along the various flats before mk08 was displaced up along the footwall ramp on mk07. there should therefore be a stepping down of the trailingend sheet in the zone. this would correspond to translation along the lower flat level of the mk09–mk10 thrust sheets, which facilitated the formation of a depression above mk10, where the rubjerg knude formation was deposited and preserved in the most proximal piggyback basin of the glaciotectonic complex. the leading zone is characterised by carrying a relatively high-elevated piggyback basin, where the rubjerg knude formation was deposited on an uneven erosional unconformity. during the early phase of imbrication, this piggyback basin was separated into five sub-basins, before they were finally trapped by overthrusting and deposition ceased. the accumulated displacement in the leading zone is c. 180 m. the thrusting probably started from a detachment level in the upper flat (10 m level), inferred from the thickness of mk02–mk04. thrusting then shifted down to the second flat (20 m level). assuming that the first half of the displacement started as an imbrication of the mk02–mk06 thrust sheets, then lateral translation of the upper 10 m thrust-sheet segment resulted in lateral displacement of the upper thrust level in the order of 100 m. subsequently, the detachment surface was lowered down to the 20 m level, and it is evident that this detachment surface is the next flat level, along which about 100 m lateral translation occurred. one of the main lines of evidence that this level is another pervasive flat level is that it acted as an upper flat for the displacement of mk01. the propagation of this thrust sheet probably started with a minor dislocation along the 30 m flat level, which is known to be a pervasive flat level from the moserende section, before it moved down to be a dislocation along the lower décollement level (40 m flat level). from the lower décollement level, mk01 ramped up to the 20 m flat level along which translation occurred over a distance of 80 m before its hanging-wall flat and ramp was ramped up to the surface along the footwall ramp at the trailing ramp of the moserende section. during displacement, the mk01 thrust sheet carried the mk02–mk07 sheets piggyback resulting in over-steepening of these thrust sheets towards the trailing end (mk05 and mk07). ribjerg section the northern termination of the rubjerg knude glaciotectonic complex is the sandy hill at lønstrup called ribjerg. most of the coastal clif f below ribjerg is now protected, and vegetation covers the cliff exposures at ribjerg. however, on the south-western side of ribjerg a funnel-shaped gully has been formed by steady erosion due to high groundwater drainage in the glaciofluvial sand (fig. 111). at the boundary between the sand and the underlying mud, groundwater wells up and creates quicksand. thus, although a section through the glaciofluvial sand is well exposed, access is difficult and potentially dangerous. in spite of such obstacles, a detailed log of the succession has been measured, and the locality yields the type section of the ribjerg formation. in addition, the section is the site for studying the glaciotectonic unconformity above the skærumhede group, cropping out at the ‘lille blå’ (northernmost part of the cross-section in plate 1). ‘store blå’ and ‘lille blå’ north of the mårup kirke section, the unconformity above the mud-rich lønstrup klint formation dips gently to the north. jessen (1918, 1931) named this part of the cliff ‘det store blå’ and ‘det lille blå’ (the 145 big blue and the small blue, respectively, a reference to the blue colour of the clayey mud in the mud-rich part of the cliff section). in general, the mud is a mobilised succession with only few bedding surfaces and thrust faults preserved. in the store blå cliff section, the structural features recorded accord well with the maximum compressional model described for the duplex units of the mårup kirke section (fig. 110). in the lille blå cliff section, the mud is structureless, and no primary bedding surfaces are preserved. a secondary sub-horizontal planar fabric is recognisable, and pebbles and boulders occur on the unconformity as well as in the uppermost metre just below the unconformity. jessen (1931) interpreted the lille blå as dislocated older yoldia clay, which he named portlandia arctica clay after the occurrence of the identified mollusc species in the unit. jessen’s description of the clay compares well with the characterisation of the skærumhede group and the model of glaciodynamic development presented below. due to the progressive deformation in the proximal part of the glaciotectonic complex, deeper levels of the skærumhede group were thrust up into a position close to the main l/r-unconformity level, such that an increasing proportion of the group has been eroded. jessen (1931) also described another important feature related to the lille blå cliff section. before 1895, it could be observed that the unconformity was folded into a syncline with a fold axis directed n–s. this is of course unusual since all structures described until now are assumed to have been formed by compression directed n–s due to the advance of the ice cap from the north, resulting in mainly e–w-trending structural features. the n–s-orientated fold axis is interpreted to be related to deformation by ice advance from the east, an event that also deposited the mid danish till formation. tectonic architecture the ribjerg section is defined as the section between the northern boundary of the mårup kirke section and the end of the lønstrup klint cliff section, which terminates at the vegetation-covered cliffs below the town of lønstrup. the southern boundary of the section is situated where four unconformities are superfig. 111. the ribjerg section viewed towards the north. the sandy cliff in the centre of the figure is the type locality of the ribjerg formation. photograph: july 1994. 146 imposed upon each other. these are: (1) the l/r-unconformity, (2) the glaciotectonic unconformity below the kattegat till formation, (3) the unconformity between the kattegat till formation and the glaciodynamic succession related to the ne-ice advance, and finally (4) the unconformity between the glaciodynamic successions and the vendsyssel formation (plate 1). the first three unconformities are here collectively termed the blå-unconformity. the blå-unconformity dips at 2–3° to the north. the surface is relatively planar, but uneven. a few clasts remain in depressions on the surface, but clasts protruding into the surface from below are more common. the unconformity between the ribjerg formation and the vendsyssel formation is an erosional surface dipping gently to the south. the main lithology in the vendsyssel formation is the saxicava sand, which consists of sandy heteroliths. these beds onlap the unconformity, which probably was subaerially exposed before inundation by the rising younger yoldia sea. sedimentary units in the ribjerg section, four sedimentary units are represented: the skærumhede group, the ribjerg formation, the mid danish till formation and the vendsyssel formation (figs 14, 17, 33). skærumhede group the skærumhede group comprises two formations: the stortorn formation and the lønstrup klint formation. in the southernmost part of the section (at the store blå), it is possible to distinguish the two formations (fig. 17). however, in the northern part of the ribjerg section, pervasive mobilisation has obliterated the primary lithological differences and the sediments may only be referred, undifferentiated, to the skærumhede group. in the southern part of the section, the stortorn formation constitutes the lowermost 10 m of the cliff section (fig. 17). here a cataclastic breccia separates the stortorn formation from the lønstrup klint formation above. it is inferred that this breccia represents one of the thrust-fault flats that form the boundary of the duplex segments building up the duplex units of the section. the skærumhede group is truncated by the blåunconformity, above which the vendsyssel formation was deposited. blå-unconformity the blå-unconformity is considered to represent three superimposed unconformities. the first one is the l/runconformity, the existence of which is only rarely demonstrable in this section. the second unconformity is the glaciotectonic unconformity below the kattegat till formation. the kattegat till formation has been almost completely eroded away from the ribjerg section, but is present in small, isolated pockets (fig. 31). however, the glacitectonite related to the subglacial deformation below the kattegat till formation is well preserved in a zone more than 1 m thick below the blå-unconformity (fig. 32). erratic clasts are common in this zone, probably lodged into the soft sediment from the till above, and an indicator boulder of larvikite has been recognised. a number of clast fabrics have been measured, which show a n–s long-axis orientation (variation from 010° to 175°). the unconformity is preserved at the base of the vendsyssel formation in the northern part of the mårup kirke section (fig. 37). the third unconformity is the erosional surface upon which the ribjerg formation was deposited. the creation of this surface removed much of the evidence of the preceding unconformities; indeed, at the southern extent of the unconformity, the ribjerg formation is also absent, and the vendsyssel formation rests on the composite surface. ribjerg formation the c. 25 m thick glaciofluvial sand of the ribjerg formation dominates the ribjerg section (fig. 33, plate 1). the formation comprises fineto medium-grained sand, coarsening upwards into gravel-dominated beds at the top (fig. 33). the formation was deposited on the erosional unconformity capping the skærumhede group (the blå-unconformity). at this surface, a residual coarse clastic bed is present, less than half a metre in thickness, dominated by clayey clasts derived from the unit below. the clayey clasts continue to appear in the sand beds in the lowermost 5 m of the formation. the middle part of the formation is dominated by trough cross-bedding, and the flow direction indicat147 ed from measurements of foreset beds was from east to west. the fill of the large channels incised into the medium-grained sand package also include gravel and slumped diamictite material. water-escape dykes and sand-filled cracks are common in the sand within the large channels (fig. 34). the formation coarsens upwards into a trough cross-bedded sandy gravel in the uppermost 3 m, just below the diamictite referred to the mid danish till formation. mid danish till formation the mid danish till formation is a c. 3 m thick unit of grey brown to light yellowish brown sandy till that overlies the ribjerg formation (figs 14, 33, 35). the till is divided into lower and upper beds. the lower bed is a laminated to thin-bedded, fine-grained sandy, matrix-supported diamict. lamination and bedding is deformed into irregular intraformational slump folds with fold axes trending n–s, indicating a slump-slide direction towards the west, and the unit is interpreted as a sediment gravity flow or flow till (dreimanis 1988). the upper bed is a massive, structureless and sandy matrix-supported diamict (fig. 35). the clasts, pebble to cobble in size, occur randomly, and the till fabric shows an a-axis orientation gently dipping towards the east. the unit is interpreted as a basal lodgement till (dreimanis 1988) superposed on the flow till and deposited by an ice stream moving from east to west. the mid danish till formation is truncated by the erosional unconformity upon which the vendsyssel formation was deposited. vendsyssel formation in the ribjerg section, the vendsyssel formation truncates the mid danish till formation, the ribjerg formation and the blå-unconformity. the maximum thickness in this part of the lønstrup klint section is about 12 m, decreasing towards the north, where it onlaps the unconformity above the ribjerg and mid danish till formations (fig. 33). the vendsyssel formation comprises laminated mud and thin-bedded finegrained sandy heteroliths, which in the southern part of the ribjerg section are characterised by well-preserved trace fossils created by the bivalve hiatella arctica, often with the shells preserved in life position (fig. 41). structures in the ribjerg section, the most important structures are the anastomosing joints related to the glacitectonite below the blå-unconformity (fig. 32). at the lille blå locality, the rhomb-shaped segments, 0.5–3 m in size, bounded by conjugate shear joints, are flatlying. the angle between conjugate joints varies from 10–30° and the zone-axis is orientated more or less e–w. at the store blå locality, the shear joints are more parallel with a spacing of c. 30 cm between the almost horizontal fractures, and in the southernmost part of the section, sand-fill intruded the fractures to create rhomb-shaped segments in a sandy mud matrix. interpretation of glacial geology and stratigraphic development in the interpretation presented here, the blå-unconformity is considered to be a modulation surface or deformational layer below the advancing front of the norwegian ice. the unconformity may even be interpreted as the surface onto which the sole of the ice pressed during the propagation towards the glaciotectonic complex developing in front of it. after the ice retreated, a hill-and-hole pair formed. rubjerg knude is here viewed as the hill and the depression extending to the north of the northward-dipping unconformity corresponds to the hole. the hole was subsequently filled with glaciofluvial sands (the ribjerg formation) that are younger than the rubjerg knude formation. on top of the ribjerg formation, jessen (1931) described a sandy till that is here referred to the mid danish till formation, but he also recorded a single till-bed intercalated in the meltwater sand. this sandy till as well as the thin diamictite layers related to the slumps in the troughs and channels are interpreted as precursors to the flow till that initiated deposition of the mid danish till formation. the ribjerg and mid danish till formations were formed as proglacial and subglacial units during the advance of the ice from the east towards the west with a source area in central sweden. this ice advance was also responsible for the gentle folding of the blå-unconformity and the beds above it into a syncline with a n–s-trending axis, as noted by jessen (1931). geological survey of denmark and greenland bulletin 7, 2004, p 25-28 25 recent danish seismological projects involving neotectonic investigations and structural studies have determined the edge of the baltic shield underlying denmark. the most active earthquake zones in denmark are located in northwestern jylland and adjoining offshore areas, and in the region around kattegat, øresund and north-east sjælland (fig. 1). this pattern was originally recognised by lehmann (1956) and has been confirmed by several later studies, e.g. gregersen et al. (1998). recent, more detailed investigations have documented that changes in the pattern of earthquake activity have occurred within a short time span. the most pronounced example of change – possibly related to exploitation of hydrocarbons – is an activity recorded in the central graben area of the north sea that was first documented by gregersen et al. (1998). the south-western margin of the precambrian baltic shield separates areas of different earthquake activity (fig. 1; gregersen et al. 1991). although lithospheric stresses are more or less uniform in northern europe, there are pronounced differences in the behaviour of the lithosphere across denmark. the north-eastern area underlain by the baltic shield experiences brittle failure as recorded by common earthquakes, whereas earthquakes are virtually absent in the region southwest of the shield (fig. 1). the margin of the baltic shield as defined by earthquake activity is not identical with that distinguished structurally in sedimentary studies (eugeno-s working group 1988; vejbæk & britze 1994), in crustal studies (abramovitz & thybo 2000), or by recent studies of the structure of the subcrustal lithosphere (gregersen et al. 2002; shomali et al. 2002). the physical edge of the baltic shield cannot be uniquely determined on the basis of seismological studies. the earthquakes recorded, although of low magnitude, do give information about the released stresses. the earthquakes seem to be a response to a dominant nw–se compression, also apparent elsewhere in scandinavia and northern europe geological survey of denmark and greenland bulletin 7, 25–28 (2005) © geus, 2005 seismology: neotectonics and structure of the baltic shield søren gregersen, martin glendrup,tine b. larsen, peter voss and hans peter rasmussen 5°e 54°n 58°n 100 km denmark s j ø kattegat 5 stz rfh rfh cg rfh el north sea 10°e 15°e fig. 1. computed locations of earthquakes around denmark during the last 75 years, supplemented by locations provided by the university of bergen based on arrival time readings from at least five seismograph stations. at shallow depths the south-western boundary of the baltic shield coincides with the sorgenfrei–tornquist zone (stz). cg, central graben; el, elbe line; j, jylland; rfh, ringkøbing –fyn high; s, sjælland; ø, øresund. (slunga et al. 1984; slunga 1989; gregersen 1992; müller et al. 1992). these stresses are part of the large-scale stress systems associated with continued plate motion pattern (gregersen & basham 1989; zoback et al. 1989). in contrast to present low-magnitude earthquakes, postglacial sediments in northern scandinavia have preserved features interpreted as caused by earthquakes of magnitudes around 7; these major, c. 9000 years old earthquakes are believed to be related to the post-glacial uplift of scandinavia (e.g. arvidsson et al. 1991; gregersen 2002). earthquakes are always related to fault activity, but attempts to link recent earthquakes occurring in and around denmark to geologically known faults have only been partly successful (gregersen et al. 1996). the most significant fault zone in denmark, the sorgenfrei–tornquist zone, is only locally active. recent geodetic and seismic investigations demonstrate that the two sides of the sorgenfrei–tornquist zone are characterised by different patterns of deformation, but the zone itself is not defined by a present-day seismicity trend crossing the central parts of denmark (fig. 1). seismological monitoring in denmark seismological monitoring in denmark has until recently been carried out by a staff of four full-time workers and one temporary employee at the danish national survey and cadastre (kms). this function was transferred to the geological survey of denmark and greenland (geus) in 2004. the monitoring is carried out with four permanent seismographs located in denmark, and four permanent and ten temporary stations located in greenland. currently a technical transfer from analogue to digital data handling is almost complete. this involves changes of the data processing for local earthquake signals and substitution of the earthquake location program by an improved version described below. the earthquake list covering the danish area has recently been updated to include all recognised small earthquakes (see www.geus.dk). supplementary data on earthquakes, which other agencies have located in the danish area, can be found on the web pages of the university of bergen in norway (www.ifjf.uib.no/seismologi) and the university of helsinki in finland (www.seismo.helsinki.fi). recomputation of earthquake locations the computer program hitherto used in denmark for computing earthquake locations (hypolg) is a modified version of the globally distributed location program (hypo71) from the united states geological survey (lee & lahr 1972; gregersen 1979). the current location program (hypocenter) now used at geus is that of lienert et al. (1986) and forms part of the analysis program for digital seismograph data (seisan) from the university of bergen (havskov & ottemöller 1999). the consistency of new and old procedures has been tested by analysis and comparison of all the seismogram records of danish earthquakes for the period 1995–2002 (fig. 2). the input to the two programs is the same. the hypolg program has an option for disregarding those readings that are not in agreement with the rest, i.e. have arrival time anomalies larger than three standard errors, while such anomalous readings are preserved and included in the computations in the seisan version of hypocenter. this difference has been taken into consideration in the comparison of the two programs, such that the same readings were included in both. the distance between any two computed locations of the same earthquake is typically 2 to 3 km. the largest differences of 7 to 13 km determined for four earthquakes are due either to few station readings (4–5) or poor azimuth coverage (overweight of readings from a small azimuth interval). the error in location determination for the same earthquake decreases with increased input of more arrival times. 26 denmark 57°n 55°n 100 km 4°e 8°e 12°e 16°e fig. 2. comparison of computed earthquake locations for the period 1995–2002, using the hypolg software (yellow dots) and the new seisan software (red dots). the results are very satisfactory and the shift to the earthquake location program hypocenter included in seisan (lienert et al. 1986; lienert & havskov 1995) is made for 2003 and the following years. the errors in the locations measured through the root mean square of the travel time anomalies with seisan are no higher than 3.1 sec. with an average of 1.76 sec. hypolg has a root mean square average of 1.88 sec. for the location of the 76 earthquakes. tor project a major international teleseismic project implemented in 1996 had the objective of delineating the lithospheric differences and the shape of the edge of the baltic shield across the sorgenfrei–tornquist zone. the tor project (teleseismic tomography across the tornquist zone) represents a collaboration between geophysical and geological institutes in sweden, germany, poland, france, the czech republic, switzerland, the netherlands and the usa with danish leadership (gregersen et al. 2002). field work and interpretation studies carried out by most partner institutions were completed in the years 1996–2002, except for a danish ph.d. study that is still active. the tor seismograph array (fig. 3) covered areas where extensive seismic crustal studies (i.e. in the depth range 0–60 km) over the last two decades have identified a terrane boundary between the ancient plates of baltica and avalonia. the tor project extends below this level into the subcrustal parts of the lithosphere (i.e. depth range 50–300 km). methods of investigation have included studies of pand s-wave travel time tomography, surface wave dispersion, receiver function, sks splitting and scattering analyses (shomali et al. 2002). in the tomographic cross-section (fig. 4) the local deviations from a homogeneously layered reference model are shown. sharp and steep boundaries between ‘fast’ and ‘slow’ 27 58°n 4°e 12°e 20°e 54°n 50°n 200 km denmark norway fig.4 sweden poland germany the netherlands czech republic fig. 3. the locations of seismometers during the field work for the tor project 1996–1997. red dots, short period seismographs; blue dots, broad-band seismographs; green line, deep crustal to mantle crosssection shown in fig. 4. 0 3 2 1 0 –1 –2 –3 100 200 300 51 52 53 54 55 56 57 58 °n el rfh stz avalonia crust baltica crust a b c d p -v el o ci ty p er tu rb at io n ( % ) d ep th ( km ) fig. 4. deviations in p-wave velocities in a deep crustal to mantle cross-section acquired to a depth of 350 km during the tor project (from shomali et al. 2002). for location of line see fig. 3. the colouring indicates velocities lower (red) or higher (blue) than the average at any level. the edge of the baltic shield as defined by sedimentary basin studies is located by the sorgenfrei–tornquist zone (stz). however, crustal transition occurs gradually within frame a between the ringkøbing–fyn high (rfh) and the elbe line (el). a subcrustal transition occurs within frame b, and another pronounced and steep transition occurs in frame c, roughly coinciding with the stz. a third, deep and less pronounced lithospheric transition occurs beneath the baltic shield in frame d. 28 lithosphere are seen as changes from red to blue in frames b and c, and the change from light blue to dark blue (frame d). together these boundaries represent a major change in the lithosphere between the precambrian shield in scandinavia and areas dominated by phanerozoic deposits in central europe. although temperatures and pressures are large at the investigated depths, it should be noted that the changes are stepwise and not gra-dual, as might be expected. the disparity in the position of the edge of the shield as defined at shallow crustal levels in the flanking sedimentary basins by the sorgenfrei–tornquist zone, and that at deep crustal and subcrustal levels is significant. recent crustal studies have indicated that the transition occurs somewhere between the elbe line and the ringkøbing–fyn high (fig. 4, frame a; abramovitz & thybo 2000). the subcrustal lithosphere edges revealed by the tor investigations (in frames b, c and d of fig. 4) appear to be related to the shallow crustal transitions. however, the actual relationships between the changes observed at shallow and deep crustal levels are not well understood, and further geophysical studies are desirable. references abramovitz, t. & thybo, h. 2000: seismic images of caledonian lithosphere-scale collision structures in the southeastern north sea along mona lisa profile 2. tectonophysics 317, 27–54. arvidsson, r., gregersen, s., kulhanek, o. & wahlström, r. 1991: recent kattegat earthquakes – evidence of active intraplate tectonics in southern scandinavia. physics of the earth and planetary interiors 67, 275–287. eugeno-s working group 1988: crustal structure and tectonic evolution of the transition between the baltic shield and the north german caledonides (the eugeno-s project). tectonophysics 150, 253–348. gregersen, s. 1979: earthquakes in the skagerrak recorded at small distances. bulletin of the geological society of denmark 28, 5–9. gregersen, s. 1992: crustal stress regime in fennoscandia from focal mechanisms. journal of geophysical research 97, 11821–11827. gregersen, s. 2002: earthquakes and change of stress since the ice age in scandinavia. bulletin of the geological society of denmark 49, 73–78. gregersen, s. & basham, p.w. (eds) 1989: earthquakes at north-atlantic passive margins: neotectonics and postglacial rebound, 716 pp. dordrecht: kluwer academic press. gregersen, s., korhonen, h. & husebye, e.s. 1991: fennoscandian dynamics: presentday earthquake activity. tectonophysics 189, 333–334. gregersen, s., leth, j., lind, g. & lykke-andersen, h. 1996: earthquake activity and its relationship with geologically recent motion in denmark. tectonophysics 257, 265–273. gregersen, s., hjelme, j. & hjortenberg, e. 1998: earthquakes in denmark. bulletin of the geological society of denmark 44, 115–127. gregersen, s., voss, p., shomali, z.h. & tor working group 2002: summary of project tor: delineation of a stepwise, sharp, deep lithosphere transition across germany–denmark–sweden. tectonophysics 360, 61–73. havskov, j. & ottemöller, l. 1999: seisan earthquake analysis software. seismological research letters 70, 532–534. lee, w.h.k. & lahr, j.c. 1972: hypo71: a computer program for determining hypocenter, magnitude, and first motion pattern of local earthquakes. u.s. geological survey open file report 72-224, 100 pp. lehmann, i. 1956: earthquakes in denmark. bulletin of the geological society of denmark 13, 88–103 (in danish with abstract in english). lienert, b.r. & havskov, j. 1995: a computer program for locating earthquakes both locally and globally. seismological research letters 66(5), 26–36. lienert, b.r., berg, e. & frazer, l.n. 1986: hypocenter: an earthquake location method using centered, scaled, and adaptively damped least squares. bulletin of the seismological society of america 76, 771–783. müller, b., zoback, m.l., fuchs, k., mastin, l., gregersen, s., pavoni, n., stephansson, o. & ljunggren, c. 1992: regional patterns of tectonic stress in europe. journal of geophysical research 97, 11783–11803. shomali, z.h., roberts, r.g. & tor working group 2002: non-linear body wave teleseismic tomography along the tor array. geophysical journal international 148, 562–574. slunga, r.s. 1989: focal mechanisms and crustal stresses in the baltic shield. in: gregersen, s. & basham, p.w. (eds): earthquakes at northatlantic passive margins: neotectonics and postglacial rebound, 261 –276. dordrecht: kluwer academic press. slunga, r., norrman, p. & glans, a.-c. 1984: seismicity of southern sweden, 106 pp. stockholm: forsvarets forskningsanstalt. vejbæk, o.v. & britze, p. (compilers) 1994: geological map of denmark, 1:750 000. top pre-zechstein (two-way travel time and depth). danmarks geologiske undersøgelse kortserie 45, 8 pp. zoback, m.l. et al. 1989: global patterns of tectonic stress. nature 341, 291–298. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: sg@geus.dk geological survey of denmark and greenland bulletin 13, 2007, 73-76 a series of cenozoic basins fringes the vietnamese coastal margin, often characterised by more than 10 km of sedimentary infill (fig. 1). greater parts of the margin are still in an early explorational state, although significant petroleum production has taken place in all but the southern song hong and the phu khanh basins. this has increased the need for a fundamental understanding of the processes behind the formation of the basins, including analyses of potential source rocks. the basins fringing the indochina block provide excellent evidence of the geological evolution of the region, and the basin geometries reflect the collision of india and eurasia and the late cenozoic uplift of south indochina (rangin et al. 1995a; fyhn et al. in press). in addition, the basins provide evidence of regional palaeogene rifting and subsequent late palaeogene through early neogene sea-floor spreading in the south china sea. apart from the regional cenozoic tectonic record, the basins contain a high-resolution climatic record of south-east asia due to the high depositional rates, changing depositional styles and large hinterland of the basin (clift et al. 2004). background since 1995 the geological survey of denmark and greenland (geus) and the department of geography and geo logy, university of copenhagen, have operated jointly in vietnam aiming to improve the local geoscientific capa city. the work is part of the enreca project (enhancement of research capacity in developing countries), funded by the danish international development agency (danida). this part of the enreca project focuses on an assessment of the hydrocarbon potential of the vietnamese continental margin, and has led to basin evaluations of the song hong and the phu khanh basins (fig. 1), and to a series of both vietnamese and danish m.sc. projects (nielsen et al. 1999; nielsen & abatzis 2004; andersen et al. 2005; boldreel et al. 2005; fyhn et al. in press). the ongoing second phase of the project focuses both on training vietnamese m.sc. and ph.d. students and on evaluating the hydrocarbon potential of the vietnamese part of the malay and khmer basins, as well as cenozoic evolution of the vietnamese coastal margin michael b.w. fyhn, lars henrik nielsen and lars ole boldreel © geus, 2007. geological survey of denmark and greenland bulletin 13, 73–76. available at: www.geus.dk/publications/bull 73 fig. 1. map showing major cenozoic basins and oceanic crust and simplified cenozoic structural features. a: cross-section shown in fig. 3. b: cross-section shown in fig. 4. modified from fyhn et al. (in press). the mesozoic strata underneath and shoreward of these basins. sampling of source rocks and oil seeps and drilling of two 500 m deep, fully cored wells (enreca-1 and 2, fig. 1) as well as acquisition of shallow seismic data have been carried out as part of the basin evaluations (bojesen-koefoed et al. 2005; petersen et al. 2005). furthermore, a broader analysis of the structure and stratigraphy of the entire vietnamese margin is being carried out as a separate ph.d. study funded by the university of copenhagen. tectonic models the indochina block is situated immediately south-east of the eastern himalayan syntaxis. the himalayan orogeny thus had a major impact on the structural evolution of indo china, leading to major north-west–south-east crustal shortening in the north and to significant lateral movements along shear zones transecting and bordering the indochina block (fig. 1; morley 2002). some of the largest shear zones are the north-west-trending red river, mai ping and three pagodas shear zones. south-eastward displacement and rotation of indochina and adjacent areas produced a total left-lateral offset of several hundreds of kilometres along the three shear zones (hall 2002). tapponnier et al. (1982) suggested that the south china sea and its marginal basins formed due to complex pull-apart mechanisms in response to these left-lateral displacements (fig. 2a). alternatively, taylor & hayes (1983) suggested that the formation of the south china sea was a result of a southward subduction of ocean crust beneath borneo (fig. 2b). one of the major differences between the two models is that the subduction model predicts right-lateral displacement across a large part of the vietnamese margin, whereas the pull-apart model is associated with a left-lateral transform along the margin. the offshore red river shear zone the most extensive of the indochinese left-lateral shear zones is the red river shear zone that passes through south china and northern vietnam into the song hong basin. seismic studies of the almost 20 km deep song hong basin indicate that the basin formed in response to major palaeogene leftlateral offset along the seaward continuation of the red river shear zone (rangin et al. 1995a; nielsen et al. 1999; ander sen et al. 2005). recent studies show that the shear zone continues along the vietnamese coast in the phu khanh basin further south (fig. 1; fyhn et al. in press). the shear zone runs along the western boundary of the phu khanh basin 74 fig. 2. conceptual models of the two basic theories initially proposed for the formation of the south china sea (tapponier et al. 1982; taylor & hayes 1983). later studies have suggested various integrations of the two models (hall 2002; morley 2002; fyhn et al. in press). a: the pullapart model suggests rifting and subsequent sea-floor spreading as a result of a complex left-lateral pull-apart mechanism. b: the subduction model suggests rifting and subsequent sea-floor spreading as a result of the subduction of old oceanic lithosphere beneath borneo. note that both models infer a transform zone along the central and south vietnamese margin but with opposite relative sense of motion. modified from tapponier et al. (1982) and taylor & hayes (1983). fig. 3. cross-section of the northern phu khanh basin transecting the offshore continuation of the red river shear zone. timing of the deformations shows palaeogene left-lateral movement followed by moderate right-lateral inversion during the early neogene. the structural cut-off of the pre-rift sequence towards the shear zone is interpreted to be a result of the large left-lateral movement along the zone (see fig. 1 for location). forming a major rift structure filled by thick palaeogene synrift deposits (fig. 3). left-lateral transtension ended during latest oligocene time in the phu khanh basin, but was followed by earliest miocene structural inversion. this is interpreted to reflect a change from intense left-lateral trans tension to modest right-lateral movements along the seaward extension in the red river shear zone in the basin, corroborated by a study by rangin et al. (1995b) showing that left-lateral, coast-parallel wrench faults onshore have been inverted by right-lateral movements. the latest palaeogene termination of left-lateral movement along the offshore part of the red river shear zone in the phu khanh basin does not support neogene sea-floor spreading in the south china sea as a result of left-lateral pull-apart. consequently, neogene seafloor spreading cannot have been caused by left-lateral pullapart, but was probably forced by subduction of older oceanic crust beneath borneo. palaeogene rifting along the vietnamese margin was, on the other hand, greatly influenced by left-lateral transtension. the offshore three pagodas shear zone rifting in the malay and khmer basins south-west of vietnam was originally linked to left-lateral transtension across a seaward extension of the three pagodas shear zone (fig. 1; tapponnier et al. 1982). later models suggested right-lateral faulting along the fault zone as the forcing mechanism (polachan & sattayarak 1989), or a combination of forces related to the indochina extrusion and extension caused by subduction roll-back (morley 2001), or mantle plume emplacement (ngah et al. 1996). seismic structural analysis of the vietnamese part of the malay and khmer basins indicates that rifting mainly took place during the palaeogene, and was controlled by a steep, north-north-west-trending, downward steeping master fault, which is flanked by smaller north-west-trending conjugate normal faults (fig. 4). the master fault offsets the basement with up to more than 2 sec. twt and transects the entire study region striking towards the point at which the three pagodas shear zone enters the gulf of thailand. the master fault is therefore interpreted as an offshore fault strand of the three pagodas shear zone. the fault characteristics indicate palaeogene left-lateral transtension and thus support a close relation between extrusion of indochina and rifting in the two basins. depositional trends sea-floor spreading in the south china sea did not start until the middle oligocene, and palaeogene syn-rift sedimentation was therefore dominated by alluvial and lacustrine deposition. in the song hong basin a gradual marine transgression of the margin started after the onset of sea-floor spreading. during initial transgression siliciclastic deposition in estua ries and narrow marine pathways dominated larger parts of the basins, and carbonate growth took place on inundated highs. open marine conditions prevailed in most basins during neogene times as sea-floor spreading propagated to its max imum south-western extension. extensive carbonate growth took place on many intraand interbasinal highs south of and along the vietnamese margin up to c. 16°n during the neo gene, favoured by the open marine environment and climatic conditions. in contrast, sediment supply kept pace with subsidence in most parts of the malay and song hong basins, preventing long-lasting periods of open marine sedimentation. during late neogene time, central and southern indo china were thermally uplifted, thus significantly increasing the siliciclastic input to the marginal basins. the increased terrigenous sediment supply inhibited widespread carbonate growth off southern and central vietnam and resulted in the progradation of a distinct shelf slope, which has led to the present outline of the margin. source rocks one of the main risk factors regarding petroleum exploration in the vietnamese offshore basins is the presence of adequate source rock intervals. onshore data from the enreca-1 core through the song ba trough in central vietnam show, however, that thick intervals of excellent oiland gas-prone lacustrine mudstone and humic coals may develop even in 75 fig. 4. cross-section of the vietnamese part of the malay basin which transects a fault strand of the seaward continuation of the three pagodas shear zone. the main offset along the major fault occurred during palaeogene times as left-lateral transtension forced by the indentation of india into eurasia (see fig. 1 for location). small basins characterised by high sediment input. although the song ba trough is an order of magnitude smaller than the vietnamese offshore basins, seismic data in the latter show apparent depositional similarities suggesting the presence of similar high-quality source rocks in the offshore basins (nielsen et al. 2007; fyhn et al. in press). in addition, seismic facies analysis as well as oil and gas compositions indicate that other source rock types, such as neogene fluvio-deltaic coals, carbonaceous shales and fore-reef marls are present in some of the basins and thus testify to the great petroleum potential of the vietnamese margin (bojesen-koefoed et al. 2005; fyhn et al. in press). acknowledgements this study is a ph.d. project funded by the faculty of natural science at the university of copenhagen to the first author. funding to the enreca project was given by the danish ministry of foreign affairs through danida. vietnam petroleum institute (petrovietnam) is thanked for providing the seismic reflection and well data and giving permission to publish these. references andersen, c., mathiesen, a., nielsen, l.h., tiem, p.v., petersen, h.i. & diem, p.t. 2005: evaluation of petroleum systems in the northern part of the cenozoic song hong basin (gulf of tonkin), vietnam. journal of petroleum geology 28, 167–184. bojesen-koefoed, j.a., nielsen, l.h., nytoft, h.p., petersen, h.i., dau, n.t., hien, l.v., duc, n.a. & quy, n.h. 2005: geochemical characteristics of oil seepages from dam thi nai, central vietnam: implications for exploration in the offshore phu khanh basin. journal of petroleum geology 28, 3–18. boldreel, l.o. et al. 2005: the phu khanh basin – aspects of structural evolution and hydrocarbon potential. science-technology conference: 30 years vietnam petroleum industry – new challenges and opportunities, 24–25 august 2005. hanoi, vietnam: petrovietnam (cd-rom). clift, p.d., layne, g.d. & blusztajn, j. 2004: marine sedimentary evidence for monsoon strengthening, tibetian uplift and drainage evolution in east asia. in: clift, p.d. et al. (eds): continent–ocean interactions within east asian marginal seas. geophysical monograph series 149, 235–254. fyhn, m.b.w. et al. in press: geological evolution, regional perspectives and hydrocarbon potential of the northwest phu khanh basin, offshore central vietnam. marine and petroleum geology. hall, r. 2002: cenozoic geological and plate tectonic evolution of se asia and the sw pacific: computer-based reconstructions, model and animations. journal of asian earth sciences 20, 353–431. morley, c.k. 2001: combined escape tectonics and subduction rollback–back arc extension: a model for the evolution of tertiary rift basins in thailand, malaysia and laos. journal of the geological society (london) 158, 461–474. morley, c.k. 2002: a tectonic model for the tertiary evolution of strikeslip faults and rift basins in se asia. tectonophysics 347, 189–215. ngah, k., madon, m. & tjia, h.d. 1996: role of pre-tertiary fractures in formation and development of the malay and penyu basins. in: hall, r. & blundell, d. (eds): tectonic evolution of southeast asia. geological society special publication (london) 106, 281–289. nielsen, l.h. & abatzis, i. 2004: petroleum potential of sedimentary basins in vietnam: long-term geoscientific co-operation with the vietnam petroleum institute. geological survey of denmark and greenland bulletin 4, 97–100. nielsen, l.h., mathiesen, a., bidstrup, t., vejbæk, o.v., dien, p.t. & tiem, p.v. 1999: modeling the hydrocarbon generation in the cenozoic song hong basin, vietnam: a highly prospective basin. journal of asian earth sciences 17, 269–294. nielsen, l.h., petersen, h.i., thai, n.d., duc, n.a., fyhn, m.b.w., boldreel, l.o., tuan, h.a., lindstöm, s. & hien, l.v. 2007: a middle–upper miocene fluvial-lacustrine rift sequence in the song ba rift, vietnam: an analogue to oil-prone, small-scale continental rift basins. petroleum geoscience 13, 145–168. petersen, h.i., tru, v., nielsen, l.h., duc, n.a. & nytoft, h.p. 2005: source rock properties of lacustrine mudstones and coals (oligocene dong ho formation), onshore song hong basin, northern vietnam. journal of petroleum geology 28, 19–38. polachan, s. & sattayarak, n. 1989: strike-slip tectonics and the development of tertiary basins in thailand. in: thanasuthipitak, t. (ed.): proceeding of the international symposium on intermountain basins: geology and resources, 243–253. chiang mai, thailand: university press. rangin, c., klein, m., roques, d., le pichon, x. & trong, l.v. 1995a: the red river fault system in the tonkin gulf, vietnam. tectonophysics 243, 209–222. rangin, c., huchon, p., le pichon, x., bellon, h., lepvrier, c., roques, d., hoe, n.d. & quynh, p.v. 1995b: cenozoic deformation of central and south vietnam. tectonophysics 235, 179–196. tapponier, p., peltzer, g., le dain, a.y., armijo, r. & cobbold, p. 1982: propagating extrusion tectonics in asia: new insights from simple experiments with plasticine. geology 10, 611–616. taylor, b. & hayes, d.e. 1983: origin and history of the south china sea basin. in: hayes, d.e. (ed.): the tectonic and geologic evolution of southeast asian seas and islands 2. geophysical monograph series 27, 23–56. authors’ addresses m.b.w.f. & l.o.b., department of geography and geology, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: fyhn@geol.ku.dk l.h.n., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. 76 geological survey of denmark and greenland bulletin 35, 2016, 79-82 79© 2016 geus. geological survey of denmark and greenland bulletin 35, 79–82. open access: www.geus.dk/publications/bull a marked change in crustal thickness is seen at the deformation boundary between the undisturbed archaean core in the south and reworked archaean gneiss in the foreland of the nagssugtoqidian orogen in west greenland. in addition, intra-crustal boundaries can be tentatively interpreted. this is the first information on crustal structure in the area, which is known for kimberlite, carbonatite and ultramafic lamprophyre occurrences, and diamond exploration. the information is based on two summer seasons of passive seismological data – earthquakes – recorded on five broadband seismological stations placed on an almost 200 km long profile crossing the deformation boundary. the stations were installed in the remote area with solar panels and batteries. between 11 and 27 distant earthquakes were recorded on each of the five stations used for the receiver function analysis. geological background the receiver function profile (fig. 1) is located on precambrian rocks in central west greenland. it crosses over a deformation front marking the southern nagssugtioqidian front defined by the transition from undeformed, discordant dykes in the south to extensively deformed dykes in the southern nagssugtoqidian orogen in the north. between the inland ice and sukkertoppen ice cap this front is marked by a sharp change in aeromagnetic signatures (van gool et al. 2002). the structural front coincides with a metamorphic transition and marks the southernmost boundary of penetrative palaeoproterozoic reworking at amphibolite facies. there is a continuity of lithologies across the southern nagssugtoqidian front, and reverse, south-directed thrusting shows that the southern nagssugtoqidian orogen is a parautochthonous foreland belt (van gool et al. 2002). crustal structure over the nagssugtoqidian deformation front in west greenland: receiver function analysis trine dahl-jensen, peter h. voss and tine b. larsen sukkertoppen iskappe greenland nordre isortoq nagssugtoqidian orogen nagssugtoqidian orogen kangerlussuaq sisimiut surficial quaternary deposits sisimiut charnockite metasedimentary rocks anorthosite reworked orthogneiss granodioritic gneiss orthogneiss north atlantic craton nagssugtoqidian orogen amphibolite aasivik terrane (early archaean)50 km north atlantic craton sarfartoqsarfartoq defo rm at io n fro nt defo rm at io n fro nt broadband station kangaamiut sfjd sa4gsa4g sa3gsa3g sa2g sa1g sisg sfjd sisg fig. 1. geological map of the area (henriksen et al. 2009). the location of the deformation front is based on aeromagnetic data (rasmussen & van gool 2000). the stations used are indicated by a blue star. the station sfjd was not used. 8080 data and method receiver function (rf) analysis relies on recording naturally occurring earthquakes followed by the isolation of s-waves generated locally at geological boundaries under the recording station from incoming p-waves from distant earthquakes (ammon 1991). data acquisition to acquire data, five seismometer stations were installed along a profile with c. 34 km distance (fig. 1). four of the stations where located in the wilderness (fig. 2), and one in the town of sisimiut. the stations were equipped with güralp-3t broadband 3-component sensors and 24 bit sam data loggers, recording data locally. the stations recorded data from june 2006 to early september 2007. power was supplied by batteries, charged by solar panels. consequently the stations did not record in november, december and january when the lack of sunlight prohibited charging the batteries. for each station, teleseismic events were selected in several steps. lists of events over magnitude 5.0 were generated from international earthquake catalogues (usgs 2015), and useable events were selected by individual inspection of the recorded waveforms. the noise conditions varied from station to station.. in total, 41 individual earthquakes (fig. 3) were accepted for analysis on one or more of the five stations, resulting in 95 earthquake records for rf analysis. the number of accepted records varied from 11 to 27 on the individual stations. the events range from magnitude 5.3 to 8.3 and cover a large azimuth and distance range around sarfartoq. fig. 2. the seismic stations. the almost 100 % exposure of bedrock allows for all stations to be placed on hard rock. a: sa1g. the completed installation in 2006 with solar panels facing south, the sensor placed on bedrock and insulated, and just as the datalogger, protected against the elements. photo: k.k sand. b: sa2g was installed in 2006. photo: p.h.voss. c: sa4g packed in 2007 for return. photo: h. rasmussen. d: from sa3g. the seismic stations were deployed and returned by an air greenland as350. photo: h. rasmussen. c a d b 81 method for each of the records the rf was calculated (kind et al. 1995; yuan et al. 1997). first, a bandpass filter of 2–50 sec (0.5–0.02 hz) was applied. then the p phases were rotated in three dimensions to minimise early energy on the radial and transverse components. this procedure aligns one component (l) with the incoming p energy, one (q) with the sv energy and one (t) with sh energy. the energy from the p-s converted phases should then be isolated on the q component, and the t component should not contain any energy if the structure beneath the station is horizontally stratified. the obtained inclination and azimuth were compared to the theoretical values calculated using the iasp91 model (kennett & engdahl 1991), and events discarded if the inclination differed by more than around 5° and azimuth by more than 10–15°. the q component was then deconvolved with the l component (ammon 1991) to produce the 95 receiver functions (rfs). the final step is to depth-convert and back-trace each rf in 3d in the earth and project the data onto a profile following the method described in (yuan et al. 1997). results figure 4 shows the back-traced receiver functions from each station along the profile. the rfs overlap more and more with depth, providing a complete cover at c. 50 km depth for the four inland stations where the distance between stations is c. 34 km. the mohorovicic discontinuity (moho) is clearly seen, as well as some indications of intracrustal boundaries. the frequency content of the rf is low (0.5–0.02 hz) so we cannot expect to resolve structural details. the depth to moho varies from just under 40 km at the se (inland) end of the profile to just under 50 km at the nw end, which lies offshore. the change of depth to moho appears as a slope dipping 18° towards nw in the plane of the profile, starting at the location where the profile crosses the deformation front, and thus at the sharp boundary seen on the aeromagnetic data (rasmussen & van gool 2000; van gool et al. 2002). the crustal thickness at kangerlussuaq (sfjd, fig. 1) is 47 km (dahl-jensen et al. 2003), correlating well with the thickness along the profile nw of the deformation zone. south of the deformation front, the shallower moho also is in agreement with thinner crust reported in the archaean block in south greenland (dahl-jensen et al. 2003). the slope on moho starts directly underneath the location of the deformation front on the surface, indicating that the deformation front is fairly steep. the intra-crustal boundaries are poorly defined, due to the relatively large distance between the stations, but all indicate a dip towards nw, mimicking the slope on the moho. to the se of the deformation front the intracrustal converters seem weaker, possibly indicating the more undisturbed archaean rocks of the north atlantic craton. evaluation and outlook this study illustrates that recording distant earthquakes over a period of several months, along a profile of seismological sensors can be utilised to image local, large-scale structures, which are valuable in areas where active source seismic acquisition is very costly. the profiles obtained can provide information on crustal thickness and intra-crustal structure. higher resolution images than available in this study can be obtained if station spacing is smaller. rfs also fig. 3. in all, 41 events were selected for rf analysis recorded on one or several of the five stations on the profile (green triangle). the events (red dots) are scaled by magnitude, ranging from 5.3 to 8.3. 8282 allow mapping of deeper structures in the mantle. the study area is an exploration area for diamonds (tappe et al. 2011), and information on the lithosphere will contribute to the understanding of the generation of diamond-bearing formations. a next step will be to look for converted phases from the base of the lithosphere. these are expected to have a negative polarity as the velocities are expected to drop at the transition from the lithosphere into the asthenosphere. for example, in east greenland a similar profile outlined a fossil subduction zone (schiffer et al. 2014). acknowledgements seisuk at university of leicester lent us the instruments for data collection. the ministry of mineral resources, greenland, provided the funds for the installation, operation and recovery of the stations in greenland. references ammon, c.j. 1991: th e isolation of receiver effects from teleseismic p-waveforms. bulletin of the seismological society of america 81(6), 2504-2510. dahl-jensen, t., larsen, t.b., woelbern, i., bach, t., hanka, w., kind, r., gregersen, s., mosegaard, k., voss, p. & gudmundson, o. 2003: depth to moho in greenland: receiver-function analysis suggests two proterozoic blocks in greenland. earth and planetary science letters 205(3-4), 379–393. henriksen, n., higgins, a.k., kalsbeek, f. & pulvertaft, t.c.r. 2009: greenland from archaean to quaternary. descriptive text to the 1995 geological map of greenland, 1:2 500 000. 2nd edition. geological survey of denmark and greenland bulletin 18, 126 pp. kennett, b.l.n. & engdahl, e.r. 1991: traveltimes for global earthquake location and phase identification. geophysical journal international 105(2), 429–465. kind, r., kosarev, g.l. & petersen, n.v. 1995: receiver functions at the stations of the german regional seismic network (grsn). geophysical journal international 121(1), 191–202. rasmussen, t.m. & van gool, j.a. 2000: aeromagnetic survey in southern west greenland: project aeromag 1999. in: dawes, p.r. & higgins, a.k. (eds): review of greenland activities 1999. geology of greenland survey bulletin 186, 73–77. schiffer, c., balling, n., jacobsen, b.h., stephenson, r.a. & nielsen, s.b. 2014: seismological evidence for a fossil subduction zone in the east greenland caledonides. geology 42(4), 311–314. tappe, s., pearson, d.g., nowell, g., nielsen, t., milstead, p. & muehlenbachs, k. 2011: a fresh isotopic look at greenland kimberlites: cratonic mantle lithosphere imprint on deep source signal. earth and planetary science letters 305(1–2), 235–248. usgs 2015: national earthquake information center neic. http:// earthquake.usgs.gov/earthquakes/?source=sitenav: united states geological survey. van gool, j.a.m., connelly, j.n., marker, m. & mengel, f.c. 2002: the nagssugtoqidian orogen of west greenland: tectonic evolution and regional correlations from a west greenland perspective. canadian journal of earth sciences 39, 665–686. yuan, x., ni, j., kind, r., mechie, j. & sandvol, e. 1997: lithospheric and upper mantle structure of southern tibet from a seismological passive source experiment. journal of geophysical research 102(b12), 27491–27500. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark; e-mail: tdj@geus.dk moho moho ? 0 50 100 d ep th ( km ) 0 km 100 km 200 km 300 km deformation front senw sisg sa4g sa3g sa2g sa1g fig. 4. the profile is scaled 1:1 horizontally and vertically; no vertical exaggeration. red colour marks a conversion at an interface with higher impedance below the boundary (increasing velocity and density with depth). under each station the backtraced traces from the individual rf derived from different earthquakes can be seen as a cone, overlapping more with depth. the interpretation of moho and intra-crustal converters are marked with a thin black line. geological survey of denmark and greenland bulletin 20, 2010, 23–26 23 the distribution of sand in deltas depends on the delta regime: wave, fluvial or tidal-dominated delta (orton & reading 1993; bhattacharya & giosan 2003). during the early miocene, three delta complexes built out from the fenno scandian shield into the eastern north sea basin (rasmussen 2004). the oldest delta complex, which is informally named the billund delta, is located in jylland (fig. 1). this delta complex was mainly wave-dominated (rasmussen & dybkjær 2005; hansen & rasmussen 2008; rasmussen 2009a). recently, it has been demonstrated that in modern wave-dominated delta environments sand mostly accumulates on the updrift portion of the delta (fig. 2) whereas alternating mud and sand, e.g. barrier-lagoon complexes, occupy the downdrift portion of the delta system (bhattacharya & giosan 2003). the current study shows that most of the sand in the submarine part of the miocene wave-dominated billund delta (mainly lower shoreface and delta slope sand) was deposited downdrift to the delta front and thus differs from the foreshore and uppermost shoreface accumulation found in recent delta complexes. the aim of this study is to map the distribution of submarine delta sand in the billund delta complex. a detailed understanding of the distribution of delta sand, such as found in this delta complex, is crucial for developing predictive tools in sequence stratigraphy and for seismic interpretation, especially in the application of seismic attribute analysis of the subsurface. geological setting the eastern north sea basin was subject to inversion in the early miocene (rasmussen 2009b). the inversion tectonism resulted in high sediment input into the norwegian–danish basin and progradation of delta complexes. during the early miocene, the deltas built far out into the basin (rasmussen 2004) and were predominantly wave-dominated (rasmussen & dybkjær 2005; fig. 1). the middle – late miocene was characterised by deposition of marine, clayey sediments at water depths of more than 100 m (rasmussen 2009a). at the end of the late miocene and during the pliocene the shoreline prograded several times across denmark (rasmussen et al. 2008) and reached the central part of the north sea both during the latest late miocene and late pliocene. in the early miocene, a warm temperate to subtropical, humid climate prevailed (larsson-lindgreen 2009). the subtropical climate continued into the early middle miocene, but was succeeded by a marked climatic cooling in the middle middle miocene. apart from an interval in the early pliocene the climate was relatively cool during the remaining part of the neogene. during the miocene and pliocene, the region was located in the northern part of the zone of prevailing westerly winds, which led to a high wave energy regime in the eastern part of the north sea basin due to the long fetch across the north sea (galloway 2002; rasmussen et al. 2008). distribution and grain size of sand in the miocene wave-dominated billund delta, denmark erik s. rasmussen and jens bruun-petersen fig. 1. palaeogeographical reconstruction of the billund delta. the billund delta is located in the central part of jylland. 50 km 56°n 8°e 10°e sacalin island upd rif t downdrift lo ng sh ore cu rr en t black sea sãrãturile strandplain danube: sf. gheorghe 10 km n fig. 2. example of a modern wave-dominated delta. note the amalgamation of beach ridges on the updrift portion of the delta, and spits and barriers that enclose lagoons on the downdrift flank. yellow: sand. green-grey: mud. modified from bhattacharya & giosan (2003). © geus, 2010. geological survey of denmark and greenland bulletin 20, 23–26. open access: www.geus.dk/publications/bull 2424 sand distribution in the billund delta the development of the billund delta complex was studied from high-resolution seismic data and borehole data. the delta complex was deposited during a period of a high rate of sediment supply (rasmussen 2009a). based on a number of boreholes drilled in central jylland it is possible to reconstruct the billund delta system in the area. the n–s-orientated correlation panel shows a series of sand-rich lobes prograding across clay-rich successions (fig. 3a). the gamma-ray log of the sand-rich units shows a serrated pattern, with generally decreasing values upwards. the grain size is dominated by fineto medium-grained sand with the latter dominating in the upper part of the succession. coarse-grained sand and gravel occur in the uppermost part (fig. 3a). at the northernmost borehole, the upper part is characterised by coarsegrained sand overlain by a succession of alternating fineand medium-grained sand. the log pattern at this site is characterised by generally upward increasing gamma-ray values (fig. 3a). the sand was deposited in clinoforms, with a dip of 7–10° according to seismic data (hansen & rasmussen 2008; rasmussen 2009a). locally, at the top of the clinoformal package, erosive features are seen with concave-upward structures that are filled with a succession showing transparent or subparallel reflection patterns (e.g. rasmussen 2009a). the updrift part of the delta complex is represented by the løvlund and grindsted boreholes (fig. 3b). at these sites alternating sandand mud-rich successions dominate. the 10–20 m thick sand-rich part is dominated by grey, fingrained sand. mediumto coarse-grained sand occurs in the upper part. in the westernmost borehole, at grindsted, the lower part is dominated by 1–2 m thick sand-rich deposits intercalated in a mud-dominated succession. large cuttings show that a substantial part of the succession at this site consists of alternating thin sand beds and clay layers. the sand beds are normally graded. the downdrift flank of the delta complex is characterised by grey, mediumto coarse-grained sand. pebbles are common and clasts with diameters up to 2 cm have been found. the sand-rich succession is 20–50 m thick (fig. 3b), but seismic data indicate a thickness up to 75 m immediately north of the billund well (hansen & rasmussen 2008). thin mud layers have been found in a few samples, but mud is not a common lithology. depositional environment a prograding depositional system is indicated by the sedimentary succession characterised by s–swdipping clinoforms and the general coarsening-upward trend seen in the boreholes (fig. 3). the mud-dominated part of this prograding system is dominated by marine palynomorphs (dybkjær 2004). their concentration decreases upward, indicating a shallowing-upward succession with increasing terrestrial influence. the fining-upward succession found in the upper part of the hammerum well is interpreted as fluvial channel deposits (rasmussen et al. 2006; rasmussen 2009a). fig. 3. two correlation panels of the billund formation. a: n–s correlation panel showing a strike section of the billund formation. the length of the profile is c. 45 km. b: e–w correlation panel showing a cross section of the billund formation. the length of the profile is c. 20 km. gr: gamma-ray log. lod: last occurrence datum. m b.s.: m below surface. a b clay and clayey silt fine-grained sand medium-grained sand coarse-grained sand gravel lod of homotryblium spp. abundant lithology in borehole main boundaries depositional environments lod of d. phosphoritica common sand (marine) sand/gravel (continental) clay (marine) clay (lagoonal) vejle fjord fm vejle fjord fm billund fm 250 240 230 220 210 200 190 180 170 230 220 210 200 190 180 170 160 230 220 210 200 190 180 170 160 240 230 220 210 200 190 180 170 160 vandelbillundgrindsted løvlund m b.s. grm b.s. gr m b.s. gr m b.s. gr billund fm billund fm vejle fjord fm vejle fjord fm 240 230 220 210 200 190 180 170 160 150 220 210 200 190 180 170 160 150 140 130 120 180 170 160 150 140 130 120 210 200 190 180 170 160 150 140 130 billund m b.s. gr store vorslunde m b.s. gr fasterholt, klyngholt m b.s. gr hammerum m b.s. gr addit mb a b west east south north 25 a prograding system overlain by fluvial channels (hansen & rasmussen 2008; rasmussen 2009a) indicates a deltaic depositional environment. the development of spits and barrier complexes south-east of the main area of progradation implies a depositional system characterised by longshore transport of sediment (rasmussen & dybkjær 2005). the predominance of storm deposits with hummocky and swaley cross-stratification and other types of tempestites (rasmussen & dybkjær 2005) indicates a wave-dominated delta front. in the modern wave-dominated rhône delta, sediment (sand) transport to the delta platform and prodelta slope occurs in two ways (maillet et al. 2006). (1) during storms, erosion of the foreshore and upper shoreface leads to transport of sand partly to the outer delta platform and partly to the delta slope and (2) during floods, sand is transported in migrating dunes from the fluvial system towards the mouth bar. sand accumulations at the slope break of the delta platform may destabilise the area by increasing the angle of the delta slope (the equilibrium profile). slope failure may result from the steepened slope or from changes in pore-water pressure due to wave action, resulting in deposits being shed directly down the delta front as mass-flow deposits. high sediment supply to the billund delta occurred from the rivers (rasmussen 2009a). the supply was probably dominated by bedload transport as indicated by the braided channels that dominated the fluvial system feeding the delta (rasmussen et al. 2006). therefore, migration of dunes towards the delta platform was more effective than in the modern rhône delta, where man-made constructions have significantly reduced the sediment influx to the delta. direct sediment supply from the fluvial system to the billund delta complex is indicated by the occurrence of large clasts, up 2 cm, in the lower part of the delta slope. such large clasts have never been reported from miocene shoreface deposits (e.g. rasmussen & dybkjær 2005) indicating that the hydro-dynamic conditions in the shoreface zone was unfavourable for transport of clasts of that size. the central and the downdrift parts of the billund delta slope were thus dominated by sedimentation of mediumto coarse-grained sand with its source in the main river system. failure at the mouth bar, spit and downdrift beach thus sourced sediments for deposition on the delta slope (fig. 4). the deposition of graded, predominantly finegrained, sand beds in the updrift part of the delta complex indicates sedimentation from suspended sand clouds generated by storms or from diluted turbidity currents (fig. 4). the finer-grained character here reflects that the source was the stacked beach ridges from the updrift flank of the delta. the sand in this part of the delta has been effectively sorted during transport along the shoreline before deposition and is therefore finer grained. for example, the barrier and spit systems of the billund delta found 25 km south-east of the main delta consists of fineto medium-grained sand with few intercalations of gravel (rasmsussen & dybkjær 2005). the clasts of the intercalated gravel layers do not exceed 5 mm. discussion the distribution of sand in recent wave-dominated deltas is characterised by coherent sand accumulation, e.g. amalgamated beach ridges in the updrift portion of the delta complex (fig. 2; bhattacharya & giosan 2003). the downdrift erosion of coastline during storms (mainly fine sand) erosion of slope during storms (mainly coarse sand). clay is resuspended and exported towards the shelf failure due to instability of slope after a flood direct fluvial influx of sand and gravel during floods fluvial accretion of mouthbar during floods fig. 4. depositional model for a wave-dominated delta showing typical areas with erosion and deposition. at the river mouth fluvial sand accretes to the mouth bar, with high potential for slope failure. direct influx of sediment onto the delta slope may occur during floods. in the downdrift portion of the delta front wavereworked sand (relatively coarse-grained) forms spits and barriers that may be eroded during storms. on the updrift flank fineto mediumgrained sand is deposited as beach ridges which are a source of sediment supply to the delta slope. consequently, finer-grained sand is deposited in this area. 2626 flank is commonly dominated by river-borne clay and sand deposited in lagoons protected by spits and barriers. a different pattern of sand distribution is seen in the billund delta where most of the sand was deposited in a downdrift position of the main delta (fig. 4). this different depositional pattern can be explained by both the depositional environment and the geological setting. the billund delta prograded into a basin with relatively deep water (c. 100 m) and the delta front was relatively steep (c. 7–10°; hansen & rasmussen 2008). fluvially transported, coarse-grained sediments were at times shed directly down the delta slope as mass-flow deposits (fig. 4). the high-wave energy regime in the region and the high frequency of storms also enhanced sand transport downdrift of the main delta lobe and some of this was directed offshore beyond the delta slope break and deposited in deeper water (fig. 4). the distribution of submarine sand in the downdrift setting is important, because foreshore and uppermost shoreface sediments are rarely preserved in the geological record. therefore, in deltaic systems of the same character as the billund sand, with steeply dipping, 50–100 m high and asymmetric clinoforms, the reservoir sand is most likely found in the downdrift portion of the wave-dominated delta. this type of delta is best developed in a ramp setting that has undergone a tectonic phase, which resulted in a sudden increase in accommodation space, and is characterised by a high sediment supply due to the formation of a high relief in the hinterland. in such a setting, the fluvial system is dominated by bed-load transport and migration of dunes to the delta front is a common phenomenon. the proportion of river-borne sediment is also important. longshore currents and wave processes can move the fine-grained fraction downdrift and offshore and thereby lead to concentration of sand on the main delta platform. acknowledgements environment centres ribe, ringkøbing and århus are thanked for financial support. references bhattacharya, j. p. & giosan, l. 2003: wave-influenced deltas: geomorphological implications for facies reconstruction. sedimentology 50, 187–210. dybkjær, k. 2004: dinocyst stratigraphy and palynofacies studies used for refining a sequence stratigraphic model – uppermost oligocene to lower miocene, jylland, denmark. review of palaeobotany and palynology 131, 201–249. galloway, w.e. 2002: paleogeographic setting and depositional architecture of a sand-dominated shelf depositional system, miocene utsira formation, north sea. journal of sedimentary research 72, 447–490. hansen, j.p.v. & rasmussen, e.s. 2008: structural, sedimentologic, and sea-level controls on sand distribution in a steep-clinoform asymmetric wave-influenced delta: miocene billund sand, eastern danish north sea and jylland. journal of sedimentary research 78, 130–146. larsson-lindgren, l. 2009: climate and vegetation during the miocene – evidence from danish palynological assemblages. litholund theses 19, 20 pp. + 3 appendices. maillet, g.m., vella, c., berné, s., friend, p.l., amos, c.l., fleury, t.j. & normand, a. 2006: morphological changes and sedimentary processes induced by the december 2003 flood event at the present mouth of the grand rhône river (southern france). marine geology 234, 159–177. orton, g.j. & reading, h.g. 1993: variability of deltaic processes in terms of sediment supply, with particular emphasis on grain size. sedimentology, 40, 475–512 rasmussen, e.s. 2004: stratigraphy and depositional evolution of the uppermost oligocene – miocene succession in western denmark. bulletin of the geological society of denmark 51, 89–109. rasmussen, e.s., 2009a: detailed mapping of marine erosional surfaces and the geometry of clinoforms on seismic data: a tool to identify the thickest reservoir sand. basin research 21, 721–737. rasmussen, e.s. 2009b: neogene inversion of the central graben and ringkøbing–fyn high, denmark. tectonophysics 465, 84–97. rasmussen, e.s. & dybkjær, k. 2005: sequence stratigraphy of the upper oligocene – lower miocene of eastern jylland, denmark: role of structural relief and variable sediment supply in controlling sequence development. sedimentology 52, 25–63. rasmussen, e.s., dybkjær, k. & piasecki, s. 2006: neogene fluvial and nearshore marine deposits of the salten section, central jylland, denmark. bulletin of the geological society of denmark 53, 23–37. rasmussen, e.s., heilmann-clausen, c., waagstein, r. & eidvin, t. 2008: the tertiary of norden. episodes 31, 66–72. authors’ addresses e.s.r., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: esr@geus.dk j.b.-p., environment centre ribe, sorsigvej 35, dk-6760 ribe, denmark. geological survey of denmark and greenland bulletin 35, 2016, 31-34 31© 2016 geus. geological survey of denmark and greenland bulletin 35, 31–34. open access: www.geus.dk/publications/bull the pre-quaternary sediments and rocks in denmark generally have a low content of radioactive minerals and elements. uranium, thorium and radium are built into mineral structures or are, for example, adsorbed on the surface of clay minerals, fe-minerals or organic material. radon (222rn) is a radioactive noble insoluble gas with a half-life of 3.8 days. it belongs to the uranium (238u) decay chain where radon is formed from radium (226ra). when rn is formed by radioactive decay from ra, the emanation process sends part of the radon produced into the pore spaces of rocks and soils. from here, the radon can enter and accumulate in buildings. the source of the radioactive materials in danish sediments and rocks is primarily from weathered precambrian crystalline rocks from norway, sweden, finland and the danish island of bornholm. physical and chemical weathering disintegrates these rocks and rivers transport the material into the danish–norwegian and danish–polish sedimentary basins. several studies have analysed and described the radioactive content of danish sediments and crystalline rocks (e.g. damkjær & korsbech 1985, 1988; gravesen et al. 1996, 1999; gravesen & jakobsen 2010) and investigations have demonstrated a relationship between sediments and rocks and rn levels in danish buildings (andersen et al. 2001). this paper addresses the radioactive content of sediments and rocks with the highest radioactive levels in denmark and the highest recorded radon emanations: precambrian crystalline rocks on bornholm and late paleocene clays in north-western jylland (fig 1). the data were collected by gravesen et al. (1999) at the geological survey of denmark and greenland (geus) with the aim of characterising and mapping rn in danish rocks and sediments. methods and data in thisted, a 4 m long trench was dug, four shallow boreholes were drilled and older borehole data in geus’ jupiter database were studied. the trench was dug to a depth of c. 2 m into till deposits and limestone (fig. 2) and the lithology, structures and macro-pores were described and samples collected. the four 5–6 m deep boreholes were drilled close to the trench. a total of 48 samples of the tills, clay and limestone were described and collected every 30 cm. the samples from the trench each weighed more than 500 g whereas the core samples weighed 300–500 g. studies and sampling were carried out at the bornholm outcrops of granites and diabase at allinge-sandvig and tejn (fig. 1, loc. 3–4) at the northern end of the island and of the crystalline rocks at birkely (near the farm of vallensgård) to the south of almindingen (fig. 1, loc. 5). the lithology, structures, weathering, and faults and fractures pre-quaternary rocks and sediments with a high level of radioactivity in denmark peter gravesen and peter roll jakobsen 1 2 3 4 5 10°e sweden 57°n 55°n 10°e 100 km denmark sweden jylland germany 57°n 55°n bornholm s o rge n f re i–t o rnqu i s t fau l t zone fig. 1. map of denmark with the highest levels of radioactivity measured in surface rocks. 1: thisted. 2: erslev. 3: allinge. 4: tejn. 5: birkely. 2 m 1 m 0 1 m 2 m 3 m 4 m fractures limestone (danian) black clay (paleocene) sand sandy till illuvial zone topsoil fig. 2. geological section in the trench from the eastern part of thisted. 3232 were described. a total of 10 samples of 500–1000 g were collected from the three localities. the chemical and physical components of the matrix and the petrographical components of the clast material of the samples were analysed. the u concentration was measured by instrumental neutron activation analysis by activation laboratories ltd, canada. the ra content was measured with a germanium detector at the national institute of radiation protection, denmark. the rn emanation rate was determined using the closed-chamber method using zns(ag) scintillation cells at the risø national laboratory, technical university of denmark. relationship between danian and late paleocene deposits in thisted the distribution of danian – late paleocene (65.5–55.8 ma) deposits in north-western jylland in the area of mors and thisted is related to two salt diapirs, the erslev salt diapir (pedersen et al. 2013) and the thisted salt dome (hansen & håkansson 1980). sediment distribution around the two salt structures is similar at the pre-quaternary surface with maastrichtian chalk in the middle surrounded by danian limestone and with palaeogene clays bordering the limestone. the layers are generally inclined away from the centre of the salt structures. different aspects of the radioactive components in the areas were investigated by damkjær & korsbech (1985) and gravesen et al. (1996). the quaternary cover above the danian limestone is rather thin (0–2 m) in the thisted area. a topsoil layer overlies weichselian glacial sandy tills, clayey tills, and just above the limestone is a strongly calcareous till (about 60–80% caco3). these tills contain vertical fractures at 5–10 cm intervals. the late danian limestone is partly cemented and can be classified as a calcilutitic or calcarenitic limestone according to grain size. the trenches demonstrated more complex geological structures than expected (fig. 2). the top surface of the danian limestone was eroded and brecciated during the late paleocene and cut by horizontal and vertical fractures with pronounced karst features and inclined fractures filled with black structureless paleocene clay which contains selandian dinoflagellate cysts. the clay also contains limestone clasts and is probably redeposited. at the bottom of the karst structures, several clay-filled fractures reach at least 10–15 cm downwards (fig. 2). in a few shallow boreholes at thisted, more than 50 cm of black clay has been encountered indicating non-eroded remnants of the black clay, and black clay material is also enclosed in the till at some locations. the content of radioactive material in the thisted sediments is presented in table 1 (data from damkjær & korsbech 1988 and gravesen et al. 1999). the u, ra and rn values for the danian sediments are comparable with levels for these sediments in other parts of the country (damkjær & korsbech 1985) but the high value for the radon emanation of the black clay can only be compared with levels of paleocene black clay from the erslev area which has even higher values (table 1). high levels of radon emanations are also known from cambrian–ordovician alum shale: 16 atoms/kg/s, eocene diatomites: 22 atoms/kg/s, miocene u-bearing heavy sand: 52 atoms/kg/s, and miocene black clay: 38.9 atoms/kg/s (damkjær & korsbech 1985). discussion karst in danish limestone and chalk is found in many areas of northern jylland but the occurrence in the thisted trench is noteworthy due to its radioactivity. karst is formed by acidic water percolating through fractures in the limestone and dissolving parts of it. the erosion of the prequaternary surface started at the inversion and uplift of the sorgenfrei–tornquist zone at the end of cretaceous– paleocene time followed by cenozoic sub-areal erosion (stenestad 2006). the erosion had removed the late paleocene clay from most of the thisted area and redeposited part of the material in the shallow karst holes in the limestone. in a search for the origin of rn in buildings damkjær & korsbech (1988) investigated sediments from the erslev area and suggested that redeposited black clay in tills was the source of the high levels of radon. however, the results showed that the high values were in the limestone areas. the later investigation of andersen et al. (2001) demonallinge hammer granite 4.2 86.9 − (weathered) hammer granite 3.0 63.5 − tejn vang granite 3.8 66.0 − birkely almindingen granite 3.4 41–51 9.6 (weathered) almindingen granite 4.3 50.7 8.8 diabase (weathered) 4.2 104.2 39.6 clay: weathered diabase 8.5 − − locality lithology and age uranium radium radon ppm bq/kg atoms/kg/s thisted weichselian sandy till 0.8−1.9 18.7−24.7 7.5−10.4 late paleocene clay − − 42.4 danian limestone < 0.6 2.8−3.3 0.12−0.8 erslev paleocene clay 2.72−26.3 38−300 11.2−130 table 1. measured radioactive components 33 strated relatively high levels of rn in buildings sitting on limestone in thisted although the radioactivity of limestones at both localities was among the lowest in danish sediments. the present investigation from thisted based on the trench and shallow boreholes shows that the high levels of rn emanation from isolated or redeposited late paleocene black clays can probably be the source of the high rn levels in the buildings in this region. in sussex, england, high rn levels in houses are also partly caused by redeposited material on top of low-radioactive chalk (killip 2004). precambrian basement rocks: radioactivity and weathering bornholm is situated in the sorgenfrei–tornquist fault zone south of sweden (fig. 1). the precambrian basement of northern and eastern bornholm consists of granitic and gneissic rocks which contain abundant leucogranitic bodies, pegmatites and aplites, besides more than 250 mafic dykes that cut these crystalline rocks. the kampeløkke å locality at allinge consists of medium-grained hammer granite with a c. 20 cm thick crust of weathered granite (fig. 3). the hammer granite comprises 41% k-feldspar, 18% plagioclase, 33% quartz, 1% hornblende, 4% biotite, and accessory ti-magnetite, apatite, epidote, allanite and fluorite (micheelsen 1961). large crystals of black gadolinite, a ree-fe-be silicate mineral, are found in the area with abundant small red spots of fe2o3 on its crystal surfaces. the granite is cut by vertical and horizontal fractures. the content of radioactive compounds is seen in table 1. in the hammer granite the u content is between 3.0 and 4.2 ppm and the ra content is between 63.5 and 86.9 bq/kg. the møllebæk locality at tejn consists of coarse-grained vang granite with very coarse-grained pegmatites covered by 10–20 cm thick weathered granite. the granite contains vertical and horizontal fractures. the vang granite is composed of 33% k-feldspar, 22% plagioclase, 27% quartz, 5% hornblende, 6% biotite, 3% ti-magnetite, 1% titanite, 1% apatite and subordinate allanite (micheelsen 1961). the pegmatites have only a low content of dark minerals (averaging 1%) but 45–60% k-feldspar and 30–40% quartz. the content of radioactive components is comparable to that of the hammer granite. the birkely locality (1½ km north of the farm vallensgård) is a small quarry with medium-grained almindingen granite. the mineralogical composition is nearly the same as for the hammer granite. figure 4 shows a section with a strongly fractured and faulted part with several partly weathered olivine diabase dykes and weathered granite. these faults and fractures are orientated wnw–ese, the same direction as the major fault between the crystalline rocks and the younger sediments (micheelsen 1961), and the faults have slickensides demonstrating movement. the diabase dykes mainly contain olivine and pyroxene but biotite, epidote and hornblende also occur (callisen 1934). alteration of these minerals produced serpentine, chlorite and calcite. the diabase dykes are orientated nw– se as are many other diabase dykes and faults in these rocks. along its contacts with weathered granite, the diabase is altered to green clayey material of chlorite or serpertine (fig. 4). the rocks in the area are partly weathered and fe-bearing minerals are oxidised to yellow-brown, clayey iron-rich weathering products on the fracture surfaces. the content of radioactive components is shown in table 1. discussion in this study, the u content of the hammer, almindingen and vang granites was found to be below 5 ppm, but higher values (6–16 ppm) have previously been recorded (johansson et al. 2016). the olivine diabase dyke yielded a comparable u content, about 4.2 ppm. the granites and diabases have high levels of u, ra and emanations of rn, and examples of weathered granites and diabases with higher levels are found. some of these rocks are among those with the highest radon emanations known in denmark. the 380 180 160 140 120 100 80 60 40 20 0 cm kampeløkke å allinge møllebæk tejn hammer granite vang granite scree sandy and gravelly till weathered granite hammer and vang granite fracture pegmatite fig. 3. geological sections from allinge and tejn. 3434 weathered diabase yields rn emanation levels comparable to the paleocene clay from thisted. according to leaching investigations of pliler & adams (1962), granitic rocks become depleted in u during the first phase of acid chemical weathering, but during the subsequent weathering phases, the top of the weathered rocks become altered and now predominantly consist of resistant u-bearing minerals. the weathering of the ground surface and shallow fractures of the rocks is mainly due to hydrolysis by percolating surface water and fluctuating low temperatures. the mafic minerals in diabases, such as hornblende, pyroxene and biotite, can be altered to green chlorite or serpentine and eventually form a clayey material hosting the u. the crystalline rocks are the primary source of radon in buildings on northern bornholm, and especially where cellar walls consist of unweathered or weathered rocks the rn concentrations in the buildings can be high (andersen et al. 2001). conclusions the black paleocene clay on danian limestone in thisted and the crystalline rocks on bornholm are among the danish deposits with the highest contents of u and ra, and they also have the highest rn emanation rates. other sediments at this emanation level are fine-grained cambrian– ordovician alum shale, eocene diatomites and miocene black and brown clays and sand. most of these sediments, like the thisted clay deposits, have limited distribution and are found below thick quaternary layers and can be difficult to locate. acknowledgements the danish ministry of health is thanked for financial support. references andersen, c.e., ulbak, k., damkjær, a. & gravesen, p. 2001: radon i danske boliger. kortlægning af lands-, amtsog kommuneværdier, 74 pp. copenhagen: sundhedsstyrelsen. callisen, k. 1934: das grundgebirge von bornholm. geological survey of denmark, ii. række 50, 266 pp. damkjær, a. & korsbech, u. 1985: measurement of the emanation of radon-222 from danish soils. the science of the total environment 45, 343–350. damkjær, a. & korsbech, u. 1988: a search for correlation between local geology and indoor radon concentration. radiation protection dosimetry 24, 51–54. gravesen, p. & jakobsen, p.r. 2010: radon content in danish till deposits: relationship with redox conditions and age. geological survey of denmark and greenland bulletin 20, 39–42. gravesen, p., jakobsen, p.r. & kelstrup, n. 1996: radon i danske jordarter ii, undersøgelser og konklusioner. danmarks og grønlands geologiske undersøgelse rapport 1996/78, 113 pp. gravesen, p., jakobsen, p.r., kelstrup, n. & ernstsen, v. 1999: kortlægning af radon i danske jordarter 1. indsamling af grunddata. danmarks og grønlands geologiske undersøgelse rapport 1999/81, 58 pp. hansen, j.m. & håkansson, e. 1980: thistedstrukturens geologi – et “neotektonisk” skoleeksempel. dansk geologisk forening, årsskrift for 1979, 1–9. johansson, å., waight, t. andersen, t. & simonsen, s.l. 2016: geochemistry and petrogenesis of mesoproterozoic a-type granitoids from the danish island of bornholm, southern fennoscandia. lithos 244, 94–108, http://dx.doi.org/ 10.1016/j.lithos.2015.11.031 killip, i.r. 2004: radon hazard and risk in sussex, england and the factors affecting radon levels in dwellings in chalk terrain. radiation protection dosimetry 113, 99–107. micheelsen, h.i. 1961: bornholms grundfjæld. bulletin geological society of denmark 14, 308–349. pedersen, s.a.s., jakobsen, p.r., tougaard, l. & gravesen, p. 2013: geological map of denmark 1:50 000 – map sheet mors, nw denmark. geological survey of denmark and greenland bulletin 28, 29–32. pliler, r. & adams, j.a.s. 1962: the distribution of thorium and uranium in a pennsylvanian weathering profile. geochimica et cosmochimica acta 26, 1137–1146. stenestad, e. 2006: fluviokarst in the top of the maastrichtian chalk at rørdal, northern jutland, denmark. bulletin of the geological society of denmark 53, 93–110. schmidt equal area projection n = 22 10 m weathered granite granite diabase diabase, clayey fractures fault fig. 4. geological section from the birkely quarry, south of almindingen, bornholm. the orientation of faults and fractures, and diabase sheets (green) are shown in the diagram. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: pg@geus.dk geological survey of denmark and greenland bulletin 4, 2003, pp 25-28 25 in denmark the supply of drinking water is based almost solely on groundwater. during the past few decades danish groundwater monitoring has encountered numerous instances of pollution with pesticides and their metabolites (geus 2003a). as a result, some hundreds of abstraction wells out of about 8000 in general water supply have been closed. with this background, there is a particular concern for reducing the leaching of pesticides into the groundwater. in the present study an approach for identification of areas potentially prone to pesticide leaching is described. the potential risk of leaching of pesticides from agricultural areas into groundwater is minimised through a procedure of approval; however, some leaching still occurs (geus 2003b). the danish counties are therefore obliged to identify areas where there is a particular risk of pesticide leaching, and where restrictions in use of pesticides may be introduced to reduce the risk (miljøstyrelsen 2000). the geological survey of denmark and greenland (geus) and the danish institute of agricultural sciences (djf) have carried out a project, focusing on sandy agricultural areas, that attempts to establish the necessary background knowledge for identifying areas particularly prone to pesticide leaching. the project aims to distinguish vulnerable and less vulnerable areas, both locally and nationwide, in a cost-effective way. aim and approach the conditions under which pesticides leach from the ground surface into aquifers have been intensively investigated (e.g. flury 1996; worral et al. 2002). a general knowledge as to the circumstances and parameters that determine the fate of pesticides is thus already available. the objective of the project has been to establish a consistent set of data for danish conditions, which will allow spatial comparison of pesticide leaching vulnerability on the basis of parameters that are known, or suspected to, influence leaching. since mapping is the key to the identification of vulnerable areas, emphasis has been placed on generating a set of data based on soil parameters that can be easily and relatively inexpensively obtained. for this purpose field sites have been chosen within eight areas of western denmark (jutland) with different geological settings (fig. 1). for each site, measuring of geological and pedological profiles have been carried out and samples collected for analysis. to determine which areas are the most vulnerable to leaching, simulations have been carried out using the macro4.3 model (jarvis 2002). the relationships between soil hydraulic properties, pesticide sorption (sensu e.g. dubus et al. 2001) and disappearance (beulke & brown 2001), and the more readily obtainable soil parameters (here termed ‘inherent soil parameters’), have been investigated on the basis of data obtained from the field sites, pre-existing databases and the published literature. studies have aimed at determining the extent to which pesticides as a whole, or in characteristic groups, will leach under similar boundary conditions and reflect different combinations of inherent soil parameters. pesticide leaching in danish groundwater: identification of vulnerable areas erik nygaard, vibeke ernstsen, carsten s. jacobsen, ole h. jacobsen, rené k. juhler, peter van der keur, svend e. olesen, jim rasmussen, per rosenberg and henrik vosgerau fig. 1. location of the studied field sites in jutland, denmark. the soil types of the eight selected areas are also shown. geological survey of denmark and greenland bulletin 4, 25–28 (2004) © geus, 2004 26 simulations the study of vulnerability to leaching is focused on criteria that are robust with respect to climatic variation and land use. all the field sites chosen were on pig farms, with normal crop rotation, farm manure application, and the relevant pesticides were applied to the maximum permitted. the pesticides mcpa, metribuzin and glyphosat, and methyltriazinamin, a transformation product of tribenuronmethyl, have been studied in most detail. for glyphosat, sorption values and disappearance rates are so high that leaching could not be simulated. repeated model simulations (montecarlo type) for metribuzin indicated that leaching will occur at significant levels, while no leaching could be simulated for mcpa and methyltriazinamin (fig. 2). to illustrate the possible relationships, e.g. the hydraulic properties of the soil and the degree of leaching, eight different combinations of sorption and degradation properties of the topsoil and the lower soil were designed (fig. 3). these combinations represent a wide range in properties for which the occurrence of leaching can be simulated. such designed compounds are therefore suitable for further analysis. keeping the sorption and disappearance parameters constant, the influence of hydraulic properties on leaching can also be studied (see below). these simulations were undertaken for the sandy soil profiles stored in a djf nationwide database (fig. 2). the results indicate that maintaining the disappearance rate constant, the sorption properties of the topsoil will generally dictate the level of leaching (fig. 3). cross-plots of the simulation results for the designed compounds further indicate, by their good correlation, that irrespective of the combination of designed parameters, leaching will dominantly occur in the same soil profiles (fig. 4). fig. 2. normalised concentration of metribuzin at a depth of 1 m in all sandy soil profiles stored in a national database held at djf (one profile for every 50 km2). fig. 3. number of investigated profiles ranked according to potential for leaching of eight designed combinations of pesticides. the ranking array of profiles is different for the eight curves; see also fig. 4. correlation the input parameters for simulation (sorption, disappearance and hydraulic properties) are expensive to establish and therefore not viable for extensive or general mapping. consequently, using a correlation analysis, an attempt has been carried out in order to substitute these parameters by more readily available soil parameters. the key inherent soil parameters evaluated are soil grain size distribution, soil bulk density, organic carbon content, cation exchange capacity, ph, the content of oxalate and dithionit extractable iron and aluminium. indicators of biological activity have also been evaluated (arylsulphatase activity and substrate induced respiration). two methods of correlation analysis have been applied to study the possible correlation between inherent soil properties, hydraulic properties and pesticide specific parameters: neural network computing and multivariate data analysis. based on neural networking computing, soil water release curves have been estimated from hydraulic properties, and a satisfactory prediction of the soil-water release characteristics obtained (fig. 5). measured values of sorption and disappearance rate have been correlated with the key inherent soil parameters by multivariate data analysis (esbensen 2002). the sorption and disappearance values were predicted at the 85% level of variance for the investigated pesticides (fig. 6). the number of soil parameters required to obtain this correlation vary according to pesticide type. the content of organic carbon in the topsoil is generally the most important soil parameter, but other soil parameters such as ph and clay content were also important in some examples. grouping of pesticides extensive analyses have been carried out for a large group of pesticides at a few localities, in order to clarify whether the results with respect to the four intensely studied compounds can be applied to a broader group of pesticides. this has been investigated by correlation analysis between inherent soil properties and sorption for all the compounds. these investigations suggest that pesticides may be arbitrarily grouped in one or a few major groups, which have similar relationships to inherent soil parameters. the two compounds that do not correspond with the explanation (glyphosat and the metabolite 4 isopropyl aniline from isoproturon) are relatively strongly sorbing and degrading in sandy soils. 27 fig. 4. cross-plots of leaching of pairs of hypothetical pesticides in fig. 3. the plots illustrate the extent to which the hypothetical pesticides leak through the same profiles. a: low correspondence between profiles with the least leaking pesticides with high sorption in the topsoil and less in the lower soil. b: high correspondence between profiles with intermediate leaking pesticides with high sorption in the topsoil and none in the lower soil. c: high correspondence between profiles with the most leaking pesticides with low sorption in the topsoil and none in the lower soil. d: two families of relatively high correspondence, where one profile has high sorption and low disappearance in the topsoil, and the other has low sorption and high disappearance in the topsoil. fig. 5. measured versus neural-network predicted retention curves for three of the investigated soil types. 28 delineation in practice based on correlations of soil parameters and leaching, and grouping of pesticides, the values of soil and other parameters that may be useful for identifying soils where leaching may occur, have been investigated. knowing the rank of importance of the soil parameters, and possibly their critical values, criteria have been established for a two-phase approach in mapping. it is suggested that, initially, areas with little risk of leaching may be identified and delineated based on the few most descriptive soil parameters. these parameters can to some extent be obtained from existing geographical data, and the cost of establishing new data is relatively low. the remaining areas particularly prone to leaching can be mapped on the basis of soil parameters necessary for simulating and predicting the leaching of the pesticides. references beulke, s. & brown, c.d. 2001: evaluation of methods to derive pesticide degradation parameters for regulatory modelling. biology and fertility of soils 33(6), 558–564. dubus, i.g., barriuso, e. & calvet, r. 2001: sorption of weak organic acids in soils: clofencet, 2,4-d and salicylic acid. chemosphere 45,767–774. esbensen, k.h. 2002: multivariate data analysis in practice, 598 pp. oslo, norway: camo process as. flury, m. 1996: experimental evidence of transport of pesticides through field soils: a review. journal of environmental quality 25, 25–45. geus 2003a: grundvandsovervågning 2003, 110 pp. københavn, danmark: danmarks og grønlands geologiske undersøgelse. geus 2003b: the danish pesticide leaching assessment programme, monitoring results, may 1999 to june 2002. third report, 123 pp. copenhagen, denmark: geological survey of denmark and greenland, danish institute of agricultural sciences, national environmental research institute. jarvis, n.j. 2002: the macro model (version 4.3). technical description. available at: http://www.mv.slu.se/bgf/ miljøstyrelsen 2000: zonering. vejledning 3, 156 pp. københavn, danmark: miljøstyrelsen. worral, f., besien, t. & koplin, d.w. 2002: groundwater vulnerability: interactions of chemicals and site properties. the science of the total environment 299, 131–143. fig. 6. measured versus model-predicted sorption (kf) for mcpa using multivariate data analysis. authors’ addresses e.n., v.e., c.s.j., r.k.j., p.v.d.k., j.r., p.r. & h.v., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: en@geus.dk o.h.j. & s.e.o., danish institute of agricultural sciences, blichers allé, postbox 50, 8830 tjele, denmark. << /ascii85encodepages false /allowtransparency false /autopositionepsfiles 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<feffd5a5c0c1b41c0020c778c1c40020d488c9c8c7440020c5bbae300020c704d5740020ace0d574c0c1b3c4c7580020c774bbf8c9c0b97c0020c0acc6a9d558c5ec00200050004400460020bb38c11cb97c0020b9ccb4e4b824ba740020c7740020c124c815c7440020c0acc6a9d558c2edc2dcc624002e0020c7740020c124c815c7440020c0acc6a9d558c5ec0020b9ccb4e000200050004400460020bb38c11cb2940020004100630072006f0062006100740020bc0f002000520065006100640065007200200035002e00300020c774c0c1c5d0c11c0020c5f40020c2180020c788c2b5b2c8b2e4002e> /chs <feff4f7f75288fd94e9b8bbe7f6e521b5efa76840020005000440046002065876863ff0c5c065305542b66f49ad8768456fe50cf52068fa87387ff0c4ee563d09ad8625353708d2891cf30028be5002000500044004600206587686353ef4ee54f7f752800200020004100630072006f00620061007400204e0e002000520065006100640065007200200035002e00300020548c66f49ad87248672c62535f003002> /cht <feff4f7f752890194e9b8a2d5b9a5efa7acb76840020005000440046002065874ef65305542b8f039ad876845f7150cf89e367905ea6ff0c4fbf65bc63d066075217537054c18cea3002005000440046002065874ef653ef4ee54f7f75280020004100630072006f0062006100740020548c002000520065006100640065007200200035002e0030002053ca66f465b07248672c4f86958b555f3002> >> >> setdistillerparams << /hwresolution [2400 2400] /pagesize [595.000 842.000] >> setpagedevice geological survey of denmark and greenland bulletin 20, 2010, 63–66 63 gold exploration in the godthåbsfjord region has been carried out since the early 1990s, and the region is now recognised as a gold province. several prospects have been drilled and storø is the most advanced project in the færingehavn terrane. the gold occurrence at storø is 2635 ma old according to 207pb/206pb age determinations of metamorphic zircons associated with auriferous arsenopyrite (nutman et al. 2007). qussuk is located in the akia terrane (fig. 1), separated from the færingehavn terrane in the south by the sw–ne-trending ivinnguit fault. the ivinnguit and ataneq faults are spatially associated with several hydrothermal gold occurrences. from north to south these are: isua, storø, bjørneøen, sadelø, store malene and qilanngaarsuit (fig. 1; appel et al. 2005; kolb et al. 2009). the qussuk prospect 20–25 km north of the ataneq fault is 20 km long, 2–3 km wide, and divided from north to south into the ‘swan n’, ‘swan’ and ‘plateau’ areas (fig. 1). in this study, geology, petrography and immobile element geochemistry are used to define a vector to the ore, and we will demonstrate that the qussuk gold mineralisation shows many characteristics of orogenic gold deposits. this paper is directed towards helping gold exploration to be more efficient in the qussuk area, in the larger godthåbsfjord region and elsewhere in the achaean greenstone belts of southern west greenland. geology of the qussuk area the supracrustal rocks are deformed and metamorphosed to amphibolite grade and comprise amphibolite, ultramafic rocks, aluminous gneiss and tonalite (garde 1997). deformation is characterised by upright to overturned isoclinal folds; the rocks trend nne–ssw and are steeply dipping (garde 2008). the tonalitic orthogneiss precursor intruded into the volcanic rocks dated at 3071 ± 1 ma in the qussuk area (u-pb zircon age) whereas tonalitic orthogneiss is 3060–3000 ma old (garde et al. 2000). plagioclase-rich amphibolite units containing both biotite and hornblende are possibly of pyroclastic or volcaniclastic origin as suggested by primary textures such as graded bedding, fragmental textures and fiamme structures (garde 2007; garde et al. 2007). these primary volcanic textures and the presence of calc-alkaline and tholeiitic andesites at qussuk led garde (2007) to the conclusion that these rocks represent an archaean island arc complex. a 120 m thick sequence of leuco-amphibolite, amphibolite, aluminous gneiss, biotite schist and pegmatite dykes is found in the ‘swan n’ area (fig. 2). at the contact between the aluminous gneiss and the leuco-amphibolite a 23 m thick zone consists mainly of biotite schist with 5–30 cm thick quartz veins and quartz rods. this zone is interpreted as a hydrothermal alteration zone (schlatter & christensen 2010). continuous sampling of 1–2 m long sections reveals that the gold content averages 1.24 ppm over 23 m including several 1–2 m thick layers of biotite schist with gold contents between 3.3 and 8.4 ppm (schlatter & christensen 2010). characterisation of host rocks and hydrothermal alteration of the qussuk gold occurrence, southern west greenland denis martin schlatter and rasmus christensen akia terrane a ta ne q fa ul t nuuk 50°w52°w 64°n iv in ng ui t fa ul t 25 km ? 65°n65°n færringehavn terrane ’swan n’ and ’swan’ ’plateau’ isua qussuk storø bjørneøen qilanngaarsuit godthåbsfjord store malene sadelø qôrqut granite complex granulite facies supracrustal belts (undifferentiated) anorthosite-metagabbro complexes isua greenstone belt quaternary cover internal terrane boundary palaeoproterozoic fault structural trend line terrane boundary mesoto neoarchaean granites mesoto neoarchaean late-kinematic tonalitic-granodioritic plutons mesoto neoarchaean grey orthogneiss and granitic rocks noritic intrusions (akia terrane) eoarchaean gneiss fig. 1. geological map of the godthåbsfjord region showing gold occurrences. the qussuk prospect is divided into the ‘swan n’, ‘swan’ and ‘plateau’ areas. the red boxes show the qussuk prospect. © geus, 2010. geological survey of denmark and greenland bulletin 20, 63–66. open access: www.geus.dk/publications/bull 6464 a 100 m thick sequence with amphibolite, aluminous gneiss, biotite schist, leuco-amphibolite and pegmatite dykes occurs in the ‘plateau’ area (fig. 2). several up to 0.5 m thick quartz veins occur at or close to the contact of amphibolite and leuco-amphibolite. one of these contains visible gold and is flanked by a 0.5 m thick inner zone of semi-massive to massive pyrrhotite and an outer zone of biotite-quartz alteration products. analyses of the quartz veins and the inner alteration zone yielded up to 19 ppm gold over 0.6 m (schlatter & christensen 2010). the local rock sequences at ‘plateau’ and ‘swan n’ are different from each other and no straightforward detailed correlation can be made between the two areas, although they both belong to the same qussuk-bjørneøen metavolcanic belt (garde 2007; garde et al. 2007). the gold-enriched zones in both areas occur close to lithological contacts (fig. 2) and are structurally controlled by local shear zones and quartz veins. in the ‘swan’ area no intersection with gold concentrations above 1 ppm was encountered. methods analysed drill core sections were 20–25 cm long and a quarter of the core was used. all samples were crushed in nuuk and analysed at the actlabs laboratory in ontario, canada. gold was analysed by instrumental neutron activation. a total of 46 drill core samples and eight surface samples were used for geochemical and petrographical investigations. ten thin sections (242 spots) were analysed by a jeol jxa-8200 superprobe at the department of geography and geology, university of copenhagen to determine the chemistry of the main rock phases. hydrothermal alteration and lithogeochemical results the gold mineralisation comprises quartz veins with visible gold and massive or disseminated pyrrhotite. hydrothermal alteration zones at ‘swan n’ and ‘plateau’ comprise an inner zone of pyrrhotite, chalcopyrite and gold-quartz veins and ? bh 24 structural fw bh 25 zone w ith low er au st ru ct ur al h w st ru ct ur al h w st ru ct ur al fw st ru ct ur al fw au zo ne 1 ppb 1 ppb1 ppb 1 ppb 7 ppb18 ppb 37 ppb 63 ppb 54 ppb4250 ppb1630 ppb 1 ppb 1 ppb 38 ppb 446 ppb3 ppb 10 m ? ? bh 14 bh 02 vg vg a uzo ne an d q tz ve in in g surface 1 ppb1 ppb 1 ppb 1 ppb 15 ppb 124 ppb 114 ppb 10 ppb 10 ppb 3180 ppb 1 ppb 3 ppb1 ppb 1 ppb 1 ppb 1 ppb 1 ppb 754 ppb 26 ppb 24 ppb upper au zone at surface amphibolite leuco-amphibolite schist dolerite (containing magnetite crystals) pegmatite quartz vein abundant sulphides abundant garnet sillimanite partial melting petrographical sample with microprobe analysis qtz-bt-pl gneiss ’plateau’ n s lithogeochemical sample (wr + traces + au) petrographical sample (polished ts) ’swan n’ fig. 2. borehole logs of bh 24 and bh 25 from ‘swan n’ and of bh 02 and 14 from ‘plateau’. the geology is defined from the two bore holes and from outcrops. the alteration zones are narrower in the plateau area (a few metres) and wider in the ‘swan n’ area (tens of metres), possibly because the areas represent wider and narrower shear zones and in turn focus alteration fluids differently. only the gold contents in the lithogeochemical samples are shown. 65 fig. 3. a: ti, al, and zr are commonly considered to be immobile during hydrothermal processes and so their ratios show as primary lithological variations that have not been affected by hydrothermal alteration. the chemical groups form fairly tight clusters in a diagram based on al2o3/tio2 versus zr/tio2 and show eight main chemical groups ranging from andesite to basalt (division from barrett & maclean 1994). b: a diagram based on zr and y shows that most of the qussuk samples have calc-alkaline and transitional magmatic affinity; only a few samples plot in the tholeiitic field (division from barrett & maclean 1994). c: δk2o versus δcao+δna2o for rocks of the ‘swan n’, ‘swan’ and ‘plateau’ areas. d: δfeo versus δsio2 for rocks of the swan n, swan and plateau areas. (vf = volatile-free basis, data were normalised after loss on ignition (loi); results of mass change calculations are reported in wt% change (δ) relative to the precursor rock). –10 –8 –6 –4 –2 0 2 4 6 –2 –1 0 1 2 3 4 5 –25 –15 –5 5 15 25 35 45 –5 0 5 10 15 20 25 30 35 silicification addition of feo and silica addition of feo and loss of silica silica removed addition of k2o, cao and na2o addition of k2o and loss of cao and na2o δ k 2 o ( w t% ) δcao + δna2o (wt%) δsio2 (wt%) δ fe o ( w t% ) swan n samples from the au-zone plateau samples from the au-zone swan n and plateau samples outside the au-zone samples from the swan area andesite a andesite b basalt. andesite i basalt. andesite ii basalt x basalt a basalt b basalt c basalt d dolerite a b c d 10 20 30 40 50 0 100 200 300 400 500 600 700 z r/ t io 2 andesite a andesite b basaltic andesite ibasaltic andesite ii basalt x basalt a basalt b basalt c and d 0 100 200 300 0 10 20 30 40 50 60 70 y p p m v f zr ppm vf tholeiitic transitional calc-alkaline al2o3/tio2 an outer zone of biotite, muscovite, quartz, and sulphides (fig. 2). garnet and sillimanite also occur in the outer zone (fig. 2), but it is unclear if these are primary metamorphic or hydrothermal alteration minerals. chlorite replacing biotite and actinolite replacing pyroxene are regarded as retrograde metamorphic minerals. application of immobile element methods (barrett & maclean 1994) on 50 whole-rock analyses shows that the rocks from ‘plateau’, ‘swan n’ and ‘swan’ can be classified into ten different chemical groups ranging from andesite to basalt (fig. 3a). the rocks are mainly metabasalts with transitional to calc-alkaline affinity (fig. 3a, b). a few are tholeiitic (fig. 3b). the rocks from the gold zones in the ‘swan n’ and ‘plateau’ areas are mainly basaltic andesite i of calcalkaline affinity. hydrothermal alteration can cause mass gain or loss, which in turn results in dilution or residual concentration of the immobile elements. however, these effects do not change the initial ratio between two immobile elements for a given chemical rock type. mass changes were calculated using the single precursor approach (maclean & barrett 1993). the rocks from the ‘swan n’ and ‘plateau’ areas show gains or losses in na2o and cao and gains and small losses in k2o (fig. 3c). the rocks from the ‘swan’ area show losses in cao and na2o and gains and small losses of k2o (fig 3c). most of the samples from the ‘swan n’ and ‘plateau’ areas have gained feo and sio2, and a few samples have lost sio2 and show minor losses in feo (fig. 3d) . samples from the ‘swan’ area show losses and gains of sio2 and gains or small losses of feo (fig. 3d). 6666 discussion and conclusions garde (2007) showed that the rocks of the qussuk area were formed in a volcanic-arc setting and that they were altered during a synvolcanic and epithermal hydrothermal alteration stage, and suggested that gold was introduced during this epithermal event. the presence of plagioclase in rocks from the gold zone, however, suggests that these rocks are neither strongly altered nor extremely affected by acid leaching, because such alteration would cause breakdown of plagioclase due to leaching of na2o and cao. results from mass change calculations show that several samples from the gold zones have gained na2o and cao or lost only a small part of these mobile elements (fig. 3c). the gold occurrences from the ‘swan n’ and the ‘plateau’ areas are enveloped by biotite schist, which represents enveloping and symmetric hydrothermal alteration halos around the gold mineralisation. gains of k2o (fig. 3c) suggest that biotite formed as a result of hydrothermal alteration which is well known in orogenic gold systems (eilu & groves 2001; groves et al. 2003). the occurrence of gold-rich quartz veins and biotite-quartz-sulphide-rich hydrothermal alteration to gether with mass change calculations (fig. 3c, d) suggest that the hydrothermal fluids were enriched in k2o, sio2, feo and au. the quartz veins with visible gold and the inner pyrrhotite alteration zone cross-cut the main foliation, indicating that the veining occurred later than the formation of the foliation and shows that the gold mineralisation took place during later stages of deformation and metamorphism. metamorphic minerals of slightly lower metamorphic grade replacing minerals of higher metamorphic grade indicate that hydrothermal gold mineralisation occurred during retrograde upper greenschist to lower amphibolite facies metamorphism. the present study shows that gold in the qussuk area is controlled by post-peak, metamorphic, hydrothermal quartz veins and alteration zones akin to orogenic gold systems. however, it remains unclear whether the gold was remobilised from an earlier, possibly syngenetic enrichment or introduced from an external source. in order to find more gold-mineralised systems in the qussuk area, rocks of basaltic andesite i should be identified. this rock type is a potential target if the layer is enveloped by proximal alteration (quartz veins, pyrrhotite, elevated gold) and distal alteration (biotite, muscovite, quartz, sulphides, elevated gold concentration). acknowledgements nunaminerals a/s is thanked for financial contribution to the project. references appel, p.w.u., coller, d., vincent, c., heijlen, w., moberg, e.d., polat, a., raith, j., schjøth, f., stendal, h. & thomassen, b. 2005: is there a gold province in the nuuk region? report from field work carried out in 2004. danmarks og grønlands geologiske undersøgelse rapport 2005/27, 79 pp. + 1 cd-rom. barrett, t.j. & maclean, w.h. 1994: chemostratigraphy and hydrothermal alteration in exploration for vhms deposits in greenstones and younger volcanic rocks. in: lentz, d.r. (ed.): alteration and alteration processes associated with ore-forming systems. st. john’s: geological association of canada. short course notes 11, 433–467. eilu, p. & groves, d.i. 2001: primary alteration and geochemical dispersion haloes of archaean orogenic gold deposits in the yilgarn craton: the pre-weathering scenario. geochemistry: exploration, environment, analysis 1, 183–200. garde, a.a. 1997: accretion and evolution of an archaean high-grade grey gneiss–amphibolite complex: the fiskefjord area, southern west greenland. geology of greenland survey bulletin 177, 115 pp. garde, a.a. 2007: a mid-archaean island arc complex in the eastern akia terrane, godthåbsfjord, southern west greenland. journal of the geological society (london) 164, 565–579. garde, a.a. 2008: geochemistry of mesoarchaean andesite rocks with epithermal gold mineralisation at qussuk and bjørneøen, southern west greenland. mineral resource assessment of the archaean craton 66º to 63º30´n) sw greenland, contribution no. 8. danmarks og grønlands geologiske undersøgelse rapport 2008/4, 52 pp. garde, a.a., friend, c.r.l., nutman, a.p. & marker, m. 2000: rapid maturation and stabilisation of middle archaean continental crust: the akia terrane, southern west greenland. bulletin of the geological society of denmark 47, 1–27. garde, a.a., stendal, h. & stensgaard, b.m. 2007: pre-metamorphic hydrothermal alteration with gold in a mid-archaean island arc, godthåbsfjord, west greenland. geological survey of denmark and greenland bulletin 13, 37–40. groves, d.i., goldfarb, r.j., robert, f. & hart, c.j.r. 2003: gold deposits in metamorphic belts: overview of current understanding, outstanding problems, future research, and exploration significance. economic geology 98, 1–29. kolb, j., stensgaard, b.m., schlatter, d.m. & dziggel, a. 2009: controls of hydrothermal quartz vein mineralisation and wall rock alteration between ameralik and sermilik, southern west greenland. danmarks og grønlands geologiske undersøgelse rapport 2009/25, 76 pp. + 1 dvd. maclean, w.h. & barrett t.j. 1993: lithogeochemical techniques using immobile elements. journal of geochemical exploration 48, 109-133. nutman, a.p., christiansen, o. & friend, c.r.l. 2007: 2635 ma amphibolite facies gold mineralisation near a terrane boundary (suture?) on storø, nuuk region, southern west greenland. precambrian research 159, 19–32. schlatter, d.m. & christensen r. 2010: geological, petrographical and lithogeochemical investigations on the qussuk gold mineralisation, southern west greenland. danmarks og grønlands geologiske undersøgelse rapport 2010/10, 53 pp. authors’ addresses d.m.s., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: dms@geus.dk r.c., nunaminerals a/s, postboks 790, dk-3900 nuuk, greenland. magnetic logs from the lopra-1/1a and vestmanna-1 wells, faroe islands 41 magnetic logs from the lopra-1/1a and vestmanna-1 wells, faroe islands niels abrahamsen and regin waagstein susceptibility measurements from cores (representing basalt, lapilli-tuffs and tuffs) and magnetic logs from the lopra-1/1a well are presented. the basalts fall into highand low-susceptibility groups with no overlap. the high-susceptibility basalts (seven cores) have susceptibilities between 4 and 88 ×10–3 si and consist of basalt with < 1% vesicles from thick massive units. the low-susceptibility basalts are intergranular, intersertal or hypocrystalline and contain no or very little (< 1%) visible magnetite, are generally more altered than the high-susceptibility basalts and have susceptibilities in the range from 0.6 to 1.4 × 10–3 si (seven cores). the susceptibility of ten volcaniclastites of lapilli-tuff or tuff varies from 0.4 to 3.8 × 10–3 si. the cores from the lopra-1/1a well reveal a bimodal distribution of magnetic susceptibility. low susceptibilities ranging from 0.4 to 4 are characteristic of altered basalts poor in magnetite, lapilli-tuffs and tuffs. thus single measurements of susceptibility are of little use in discriminating between these three types of rock. susceptibility logs from the lopra-1/1a well show that the variation below 3315 m distinguishes clearly between volcaniclastics (hyaloclastites) with low and fairly constant susceptibility and basalt beds of between 5 and 10 m thickness (with high susceptibility). the volcaniclastics comprise some 60–70% of the sequence between 3315 and 3515 m with the maximum continuous sediment layer being 80 m thick. a 1½ m core of solid basalt at 2381 m and sidewall cores of basalt from the lopra1/1a well have a mean susceptibility of 22.1 ± 3.5 × 10–3 si (standard deviation (σ) = 23.6, number of samples (n) = 46), while samples of hyaloclastite (lapilli-tuff and tuff ) have a mean susceptibility of 0.85 × 10–3 si (σ = 0.39, n = 17). the mean values of the rock magnetic parameters for 303 basalt plugs from the vestmanna-1 well are: qave = 13.3 ± 0.6 (σ = 11), save = 11.8 ± 0.6 × 10–3 si (σ = 11) and jave = 4.64 ± 0.25 a/m (σ = 4.4). the reversely polarised, lowermost (hidden) part of the c. 4½ km thick lower basalt formation correlates with chron c26r. the upper (exposed) part of the lower basalt formation correlates with chrons c26n, c25r and c25n and the more than 2.3 km thick middle and upper basalt formations correlate with chron c24n.3r. keywords: magnetic logging, rock magnetism, susceptibility, nrm, magnetic reversals, faroe islands, lopra, vestmanna, north atlantic __________________________________________________________________________________________________________ n.a., department of earth sciences, university of aarhus, finlandsgade 8, dk-8200 aarhus n, denmark. e-mail: abraham@geo.au.dk r.w., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. © geus, 2006. geological survey of denmark and greenland bulletin 9, 41–49. available at: www.geus.dk/publications/bull geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1941 42 information on rock magnetic properties, susceptibility and natural remanent magnetisation (nrm) may be useful for detecting changes in rock type, structure and magnetic mineral content of rocks penetrated by boreholes. magnetic polarity is also a tool potentially of use in dating. although magnetic surveying has a long history in prospecting and mining geophysics (e.g. parasnis 1979), magnetic logging (using susceptibility) in boreholes was first developed in the 1950s with new electronic types of equipment (e.g. broding et al.1952; levanto 1958; barthés et al. 1999). logging with the purpose of magnetic polarity determinations began even later (e.g. pozzi et al.1988, 1993; bouisset & augustin 1993; ito & nogi 1995). magnetic logging instruments were developed for down-hole mapping of the magnetic field as a correlationand datingtool using magnetostratigraphy. reversals recorded in a borehole may be used for dating if they can be correlated with the geomagnetic polarity time scale (gpts). the gpts was firmly established in the early 1960s by radiometric dating of reversals recorded in young volcanic sequences on land (e.g. cox et al.1963) and by relating the polarity reversals from land to marine magnetic anomalies observed over the oceans (vine & matthews 1963). this revived the idea of continental drift and supported the new paradigm of plate tectonics. in the following years the polarity scale was extended linearly backwards through mesozoic time by correlation of long sequences of marine anomalies with the shorter land-based records (heirtzler et al.1968). the present paper deals with the results of magnetic logging of the lopra-1/1a well, situated on suðuroy, the southernmost of the faroe islands (fig. 1). the magnetic logs were acquired by schlumberger ltd. in 1997 as part of an extensive logging programme run in connection with deepening of the well. the logs cover a major part of the faroes lower basalt formation (waagstein 1988, waagstein et al. 2001). the log-like results of rock magnetic properties obtained from the continuously cored vestmanna-1 well through a younger part of the faroes basalt succession (abrahamsen et al. 1984) are summarised for comparison. magnetic logging in lopra-1/1a a geological high-resolution magnetometer tool (ghmt) was run by slumberger ltd. from 3101 to 2168 m in the deepened part of the lopra-1 well and subsequently from 3519 to 2998 m in the sidetracked lopra-1a.the kickoff depth of the sidetrack is 3091 m, which means that the two log sections overlap from 3091 to 2998 m. the two logs have been combined into a single log using an arbitrary splicing point at 3000 m. the ghmt tool records two types of magnetic measurements; the magnetic susceptibility (rmags) and the total magnetic induction (magb). examples of the records obtained are shown for the whole sequence in figs 2–5. a shorter section is shown in more detail in fig. 6. the main objective of the deepening of the lopra-1 well was to drill through the basalt formations to the expected underlying sediments. the dipole–dipole sensor susceptibility measurement tool (sumt) was therefore set to the low-resolution mode. the nuclear magnetic resonance magnetometer (nrmt) was designed to measure the total magnetic induction in the borehole within a working range of only 5000 nt around a preset expected value (schlumberger ltd., personal communication 1997) which was unfortunately much less than the actual ranges of 30 000 and 70 000 nt present within the hole. another purpose of the short working ranges applied was to protect the tool electronics, which were designed for weakly magnetic sediments rather than strongly magnetic volcanic rocks. the settings for the magnetic tools were not optimal for the basalt-dominated section actually drilled, as the streymoy vestmanna-1 lopra-1/1a 61°30´ 61°30´ 6°w 6°w 7°w 7°w 62°00´ 62°00´ torshavn vágar 10 km suðuroy fig. 1. index map of the faroe islands with the positions of the lopra-1/1a and vestmanna-1 wells indicated with stars. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1942 43 m ag ne tic s us ce pt . ( 10 –6 s l) an d to ta l f ie ld ( nt ) –10 000 0 10 000 20 000 30 000 40 000 50 000 60 000 2200 2400 2600 2800 3000 3200 3400 3600 depth (m) lopra-1 46 000 48 000 50 000 52 000 54 000 2200 2400 2600 2800 3000 3200 3400 3600 depth (m) to ta l f ie ld ( nt ) lopra-1/1a fig. 2. magnetic susceptibility (× 10–6 si) and the magnetic induction total field (nt) logged in the lopra-1/1a well between depths of 2200 and 3520 m. fig. 3. total magnetic field (magnetic induction, nt) in the lopra-1/1a well. a jump in the general level of about 4000 nt is seen at 3000 m. susceptibility was mostly outside the working range of the susceptometer. because of this, the polarity of the remanent magnetisation and hence the interplay between the susceptibility and the induced magnetisation could not be deduced from these results and a reversal chronology could not be obtained from the in situ logged data. information about the remanent polarity of the rocks drilled by the lopra-1/1a and vestmanna-1 wells (fig. 1) has, however, been obtained from drilled cores. these cores were investigated by traditional palaeomagnetic laboratory techniques and the results have been reported and presented elsewhere (schönharting & abrahamsen 1984; abrahamsen et al 1984; waagstein 1988; abrahamsen 2006, this volume). abrahamsen (2006, this volume) correlated the lowermost (unexposed) part of the c. 4½ km thick lower basalt formation with chron 26r (selandian) and the upper (exposed) part of the lower basalt formation with chrons c26n, c25r and c25n (selandian and thanetian). the more than 2.3 km thick middle and upper basalt formations are correlated with chron c24n.3r (ypresian). geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1943 44 susceptibility of cores from the lopra-1/1a borehole the magnetic susceptibility of 1 conventional and 24 rotary sidewall cores drilled at regular intervals within the deepened part of the lopra-1/1a well between 2275 and 3514.5 m has been measured in the laboratory. the cores include 14 basalts, 8 lapilli-tuffs and 2 tuffs (table 1). a single plug from each sidewall core and 22 plugs from the 1.5 m long conventional core were measured. the basalts fall into highand low-susceptibility groups with no overlap. the high-susceptibility basalts are represented by seven cores with susceptibilities between 4 and 88 × 10–3 si. they consist of basalt with < 1% vesicles from thick massive units. the texture of the groundmass varies from intergranular with a few per cent matrix (mesostasis) to hyaline with almost 50% matrix. the matrix consists of cryptocrystalline quench crystals and secondary minerals replacing glass or filling interstitial voids. the groundmass of the intergranular basalts has an estimated lopra-1/1a 10 100 1000 10 000 100 000 24002200 2600 2800 3000 3200 3400 3600 depth (m) m ag ne tic s us ce pt ib ili ty ( 10 –6 s l) fig. 4. magnetic susceptibility log from the lopra-1/1a well. the solid pale curve is a 100 point moving average (likely to be biased due to saturation of the instrument). fig. 5. magnetic susceptibility from the lopra-1/1a well (logarithmic scale). the solid pale curve is a 100 point moving average (likely to be strongly biased due to saturation of the instrument). geus bulletin no 9 7 juli.pmd 07-07-2006, 15:0644 –5000 0 5000 10 000 15 000 20 000 2200 2400 2600 2800 3000 3200 3400 3600 m ag ne tic s us ce pt ib ili ty ( 10 –6 s i) depth (m) lopra-1/1a 45 content of 3–10 vol.% titanomagnetite with a maximum size between < 0.03 and 0.2 mm. the titanomagnetite in the less crystalline basalts is too fine-grained to be estimated or cannot be seen, although the presence of an opaque or dark turbid matrix suggests that it is likely to be present. susceptibilities from seven cores from the low-susceptibility basalts vary from 0.6 to 1.4 × 10–3 si. the lowsusceptibility basalts are intergranular, intersertal or hypocrystalline and contain no or very little (< 1%) visible magnetite. they are generally more altered than the high-susceptibility basalts and lose on average about 3.7 wt% volatiles on ignition compared to 1.8 wt% for the latter group (table 1). the volatiles are dominantly crystal-bound water in secondary minerals including clay, zeolites, pumpellyite and phrenite. three of the basalts are highly vesicular with 20–25% vesicles filled with secondary minerals. the susceptibility of the ten volcaniclastites of lapillituff or tuff varies from 0.4 to 3.8 × 10–3 si with an average of 1.1 × 10–3 si. the susceptibilities of the four deepest lopra-1/1a 0 4000 8000 12 000 16 000 20 000 3200 3300 3400 3500 depth (m) m ag ne tic s us ce pt ib ili ty ( 10 –6 s i) vestmanna-1 rock magnetic properties 0 1 10 100 1000 10 000 100 000 0 50 100 150 200 250 300 350 400 450 500 550 600 650 700 depth (m) a q j s q / j( 10 –2 a /m ) / s us ce pt ib ili ty ( 10 –6 s i) fig. 7. rock magnetic properties (susceptibility, nrm and q-ratio) of the vestmanna-1 well (modified from abrahamsen et al. 1984). pale curves are 5 point moving averages (logarithmic scale). the a horizon, between the lower and upper basalt formations, is marked by an arrow and the letter a at 557 m. fig. 6. magnetic susceptibility details between 3200 and 3510 m of the lopra-1/1a well. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1945 46 r oc k ty pe h ig hsu sc ep t. ba sa lt h ig hsu sc ep t. ba sa lt h ig hsu sc ep t. ba sa lt h ig hsu sc ep t. ba sa lt h ig hsu sc ep t. ba sa lt h ig hsu sc ep t. ba sa lt h ig hsu sc ep t. ba sa lt lo w -s us ce pt . b as al t lo w -s us ce pt . b as al t lo w -s us ce pt . b as al t lo w -s us ce pt . b as al t lo w -s us ce pt . b as al t lo w -s us ce pt . b as al t lo w -s us ce pt . b as al t la pi lli -t uf f la pi lli -t uf f la pi lli -t uf f la pi lli -t uf f la pi lli -t uf f la pi lli -t uf f la pi lli -t uf f la pi lli -t uf f t uf f t uf f sa m pl e id l1 -s w c5 7 l1 -c or e 1 l1 -s w c4 6 l1 -s w c3 4 l1 -s w c3 0 l1 -s w c2 5 l1 a -s w c1 5 l1 -s w c5 9 l1 -s w c4 3 l1 -s w c4 0 l1 -s w c3 9 l1 a -s w c1 6 l1 a -s w c9 l1 a -s w c4 l1 -s w c3 8 l1 -s w c3 6 l1 -s w c3 3 l1 -s w c3 1 l1 -s w c2 6 l1 a -s w c1 3 l1 a -s w c1 2 l1 a -s w c5 l1 a -s w c1 9 l1 a -s w c6 22 75 23 80 –2 38 1. 5 24 41 26 10 27 80 30 30 33 82 22 19 24 75 25 58 25 59 .8 33 28 35 00 .5 35 31 25 60 .2 25 70 26 30 26 90 29 70 34 38 34 64 .5 35 14 .5 32 33 .5 35 12 .5 1. 26 1. 72 2. 31 1. 89 1. 3 2. 33 1. 51 6. 84 4. 94 3. 16 2. 89 2. 22 2. 62 3. 13 5. 37 4. 05 4. 35 5. 28 5. 63 4. 02 3. 84 8. 43 6. 62 4. 16 g ro un dm as s te xt ur e in te rg ra nu la r in te rg ra nu la r in te rg ra nu la r cr yp to cr ys ta lli ne hy al in e in te rg r. -in te rs er ta l in te rg ra nu la r in te rg ra nu la r in te rg r. -in te rs er ta l hy po cr ys ta lli ne in te rs er ta l in te rg ra nu la r in te rg ra nu la r in te rg ra nu la r hy al in e hy po cr ys ta lli ne hy al in e hy al in e hy al in e hy po cr ys ta lli ne hy po cr ys ta lli ne hy po cr ys ta lli ne hy al in e hy al in e m ag ne tit e vo l.% 7% 3% sp ar se ab un da nt no 10 % no < 1 % no no no no sp ar se no no v es ic le s vo l.% c or e su sc ep t. x 10 –3 s i m ax . c la st si ze ( m m ) 10 10 > 1 5 14 > 2 3 12 25 30 0. 5 2 t ab le 1 . p et ro gr ap hy a nd m ag ne tic s us ce pt ib ili ty o f c or e sa m pl es fr om l op ra -1 /1 a su sc ep t.: s us ce pt ib ili ty ; i nt er gr .: in te rg ra nu la r. pe tr og ra ph ic p ar am et er s ar e ba se d on v is ua l e st im at es o f c or es a nd t hi nse ct io ns . v ol at ile s ar e de te rm in ed fr om lo ss o n ig ni tio n co rr ec te d fo r ox id at io n of ir on . c or e su sc ep tib ili ty is m ea su re d on 8 –3 2 m m lo ng s ec tio ns o f s id ew al l c or e w ith a d ia m et er o f 2 3. 2 m m a nd o n se ve ra l p lu gs fr om c on ve nt io na l c or e 1. m ea n, m in im um a nd m ax im um lo g su sc ep tib ili tie s ar e ba se d on 1 m in te rv al s of t he g h m t lo g ce nt re d at t he c or e. t he m ea n is c om pu te d as su m in g a lo gno rm al d is tr ib ut io n of m ag ne tic s us ce pt ib ili tie s. t he s ig n ‘> ’ i nd ic at es t ha t so m e su sc ep tib ili ty m ea su re m en ts e xc ee d th e sa tu ra tio n le ve l o f t he t oo l ( 16 .4 x 1 0–3 s l). t he d iff er en ce s be tw ee n co re a nd lo g va lu es p ro ba bl y pa rt ly r ef le ct t he u nc er ta in ty in d ep th s of s id ew al l-c or es ( 0. 5 m ) an d m ag ne tic lo gs ( 0– 2 m ). m ag ne tit e m ax . s iz e (m m ) > 1 6. 4 > 1 6. 4 > 1 6. 4 > 1 6. 4 > 1 6. 4 > 1 6. 4 > 1 6. 4 2. 9 13 .5 2. 1 1. 2 > 1 6. 4 0. 6 1. 2 10 .0 4. 2 > 1 6. 4 > 1 6. 4 0. 5 0. 5 5. 5 > 1 6. 4 12 .1 > 1 6. 4 13 .9 10 .5 > 1 6. 4 7. 2 0. 4 1. 7 0. 4 0. 9 0. 5 0. 8 0. 2 3. 1 16 .1 12 .4 0. 4 0. 4 0. 4 m ea n lo g su sc ep t. x 10 –3 s i m in . l og su sc ep t. x 10 – 3 s i m ax . l og su sc ep t. x 10 – 3 s i > 1 6. 4 > 1 5. 7 > 1 6. 4 > 1 5. 7 > 1 3. 5 > 1 6. 4 > 1 4. 3 1. 4 5. 5 1. 1 1. 1 > 1 6. 4 0. 6 1. 0 2. 4 3. 9 > 1 6. 3 > 1 4. 8 0. 4 0. 5 1. 5 1. 4 0. 6 0. 7 0. 6 1. 1 0. 8 0. 9 0. 6 0. 7 1. 6 1. 8 3. 8 0. 6 0. 6 0. 6 0. 7 0. 4 d ep th v ol at ile s w t% ≤ 0. 1 ≤ 0. 2 ≤ 0. 06 ≤ 0. 03 ≤ 0. 1 ≤ 0. 08 < 1 0. 5 0 0 0 0 0 20 25 20 0 0 0 5 0 2 < 1 2 < 1 2 1 + 2 88 4– 46 47 69 77 27 63 geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1946 47 volcaniclastites from between 3438 and 3514.50 m average only 0.6 × 10–3 si. the study of cores from the lopra-1/1a well thus reveals a bimodal distribution of magnetic susceptibility. high-susceptibility rocks range between 4 and 88 × 10–3 si with the great majority falling above 15 × 10–3 si. these rocks are all relatively fresh basalts from thick massive units cooled slowly enough to crystallise titanomagnetite (visible or not). the five cores from the original lopra-1 well all consist of intergranular basalts from the massive centre of thick flows (hald & waagstein 1984) with susceptibilities between 16 and 39 × 10–3 si (schönharting & abrahamsen 1984) and thus belong to the group of highsusceptibility basalts. low susceptibilities from the lopra-1/1a borehole, ranging from 0.4 to 4 × 10–3 si, are characteristic of both altered basalts poor in magnetite (0.6–1.4 × 10–3 si), lapillituffs (0.6–3.8 × 10–3 si) and tuffs (0.4–0.7 × 10–3 si). this means that single measurements of susceptibility are of little use in discriminating between these three types of rock. the lopra-1/1a magnetic log because the recorder was run in its low-resolution mode, the susceptibilities of most of the rocks through which it passed were mostly outside its recording range. nevertheless, some general lithological features may be deduced (schlumberger ltd., personal communication 1997). the upper recording limit of 16.4 × 10–3 si is close to the typical lower limit of relatively fresh, massive basalt. this means that major intervals of saturation are a good indicator of massive basalt units. this makes it possible to divide the logged interval into two parts, below and above 3315 m (cf. figs 4–6). below 3315 m the lower part displays a highly bimodal pattern with two distinct susceptibility levels. about 85% of the logged section below 3315 m shows average susceptibilities around 0.7 × 10–3 si that we interpret as hyaloclastites. basalt layers with a thickness from 2 to 6 m are clearly identifiable within the hyaloclastites showing sharp contacts and much higher susceptibility values (to above the recorded saturation limit of 16.4 × 10–3 si). low-susceptibility basalt has been cored nearby (table 1) so the high-susceptibility intervals give only a minimum estimate of the thickness of basalt present. for comparison, a 1½ m solid basalt core from 2381 m and 24 sidewall basalt cores (abrahamsen 2006, this volume) had a mean susceptibility of 22.1 × 10–3 si ± 3.5 (one standard deviation (σ) = 23.6, number of samples (n) = 46), whereas samples of hyaloclastite (tuffs and lapilli-tuff ) had a mean value of 0.85 × 10–3 si (σ = 0.39, n = 17). these results thus compare quite well with the average of the susceptibility log data. the total magnetic field (magnetic induction) shows a jump of about 4000 nt between the two partly overlapping log runs (log sections spliced at 3000 m; figs 2–3), which must be an artifact. the magnetometer record of the total field below 3315 m (fig. 3) shows the typical effect of a highly magnetised layer within a weakly magnetised formation. the induction recorded through the volcaniclastics below and above the basalt is strongly affected by the distance to the basalt. several occurrences of such basaltic layers give rise to mixed effects through the volcaniclastics. above 3315 m the total-field magnetometer was saturated above 53 500 nt during most of the first logged section between 3101 and 2168 m (figs 2–3). this value is stronger than the local earth’s magnetic field of around 50 000 nt and could indicate a large local magnetic source of unknown origin, but is more likely a tool or calibration error. in contrast, despite the very variable character of the lower section, this is not the case for the uppermost part of the lower section (below 3315 m, fig. 3). in both cases the magnetometer was preset for maximum sensitivity of values centred at the expected value of the local earth’s magnetic field strength of 50 000 nt. above 3315 m, c. 70% of the susceptibility data (figs 2, 4–6) are greater than 16.4 × 10–3 si (the saturation level of the instrument). thick intervals above saturation level are dominant above 2550 m and reflect subaerial basalt flows. between about 2550 and 3315 m, the susceptibility log is characterised by large short-scale variations. strong variability is especially observed in the interval from 2613 to 2816 m (fig. 4). the high-frequency pattern originates from a succession of hyaloclastites and minor basalt beds. the hyaloclastites consist of lapilli-tuffs, tuff-breccias, breccias and subordinate tuffs. the variability may be explained by the presence of large clasts of basalts showing high susceptibilities set in a low-susceptibility tuffaceous matrix. only a few longer intervals of low susceptibility (< 1 × 10–3 si) can be recognised above 3315 m, the thickest ones being between 2945 and 2950 m, between 2523 and 2533 and between 2484 and 2500 m. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1947 48 the other rock magnetic properties of the lopra-1/1a well, including the magnetic polarity and correlation with the gpts, have been summarised and discussed in details elsewhere (abrahamsen 2006, this volume). rock magnetic properties in the vestmanna-1 well the 660 m deep vestmanna-1 well on streymoy (fig. 1) was drilled in 1980 through the lower part of the faroes middle basalt formation and into the top of the lower basalt formation using wireline coring technique. no magnetic in-hole logging was made during or after the drilling. however, detailed magnetic laboratory investigations of sub-sampled plugs of the fully recovered core have been published (abrahamsen et al. 1984). an illustration of most of the rock magnetic information obtained (susceptibility s, nrm intensity j, and q-ratio) is shown in fig. 7 on logarithmic scales (modified from the original data presented by abrahamsen et al. 1984). the thick pale curves are five point moving averages. the well reached 101 m into the upper part of the c. 4½ km thick lower basalt formation (waagstein 1988) whose top is indicated in fig. 7, where a 0.7 m thick basaltic conglomerate of local origin separates the lower and the middle basalt formations (waagstein & hald 1984). rocks penetrated by the overlying part of the well (0–557 m in fig. 7) all belong to the middle basalt formation. the conglomerate is stratigraphically equivalent to a c. 10 m thick sediment sequence in the south and southwestern parts of the faroe islands that includes thin beds of coal indicating a long quiescence in the magmatic activity between eruption of the lower and middle basalt formations. all three rock magnetic parameters vary more than one order of magnitude, which is not uncommon for the magnetic properties of volcanic rocks. the mean values for each of them (n = 303 samples) are qave = 13.3 ± 0.6 (σ = 11), save = 11.8 ± 0.6 × 10–3 si (σ = 11) and jave = 4.64 ± 0.25 a/m (σ = 4.4). the only readily apparent systematic trend appears to be a decrease in the susceptibility from high values at 610 m to low values at 440 m. at shallower depths, the susceptibility fluctuates around 10–2 si. the trend is mirrored in the q-ratio below 470 m, but with a slight decrease in q above this level, whereas no systematic trends appear visible in the nrm intensity. the magnetic polarity of the vestmanna-1 well was determined in detail by palaeomagnetic investigations of 303 up-oriented plugs from the fully cored borehole, that indicate a short normal polarity interval between 660 and 640 m only, all the younger samples (n = 275) being reversed (abrahamsen et al. 1984). conclusions due to instrument problems, the valuable information from the magnetic logs of the lopra-1/1a well is limited. based upon the logged susceptibility, the variation below 3315 m is clearly diagnostic between volcaniclastics (with low and fairly constant susceptibility) and basalt flows of between 5 and 10 m in thickness (with high susceptibility). between 3315 and 3515 m the volcaniclastics comprise some 60–70% of the sequence, the maximum continuous layer being 80 m thick. a 1½ m long core of solid basalt from 2381 m and sidewall cores of basalt from the lopra-1/1a well have a mean susceptibility of 22.1 × 10–3 si ± 3.5 (σ = 23.6, n = 46), while samples of volcaniclastics (lapilli-tuff and tuff ) have a mean value of 0.85 × 10–3 si (σ = 0.39, n = 17). the mean values of rock magnetic parameters for 303 basalt plugs from the vestmanna-1 well are: qave = 13.3 ± 0.6 (σ = 11), save = 11.8 ± 0.6 × 10–3 si (σ = 11) and jave = 4.64 ± 0.25 a/m (σ = 4.4). the reversely polarised, lowermost (hidden) part of the c. 4½ km thick lower basalt formation correlates with chron c26r. the upper (exposed) part of the lower basalt formation correlates with chrons c26n, c25r and c25n and the more than 2.3 km thick middle and upper basalt formations correlate with chron c24n.3r. references abrahamsen, n. 2006: palaeomagnetic results from the lopra1/1a re-entry well, faroe islands. geological survey of denmark and greenland bulletin 9, 51–65 (this volume). abrahamsen, n., schönharting, g. & heinesen, m. 1984: palaeomagnetism of the vestmanna core and magnetic age and evolution of the faeroe islands. in: berthelsen, o., noe-nygaard, a. & rasmussen, j. (eds): the deep drilling project 1980–81 in the faeroe islands. annales societatis scientiarum faeroensis, supplementum ix, 93–108. tórshavn: føroya fróðskaparfelag. barthés,v., pozzi, j.p., vibert-charbonnel, p., thibal, j. & mélières, m.a. 1999: high-resolution chronostratigraphy from downhole susceptibility logging tuned palaeoclimatical orbital frequencies. earth and planetary science letters 165, 97–116. bouisset, p.m. & augustin, a.m. 1993: borehole magnetostratigraphy, absolute age dating and correlation of sedimentary rocks, with examples from the paris basin, france. american association of petroleum geologists bulletin 77, 569–587. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1948 49 broding, r.a., zimmermann, c.w., sommers, e.v., wilhelm, e.s. & stripling, a.a. 1952: magnetic well logging. geophysics 17, 1–26. cox, a., doell, r. & dalrymple, g.b. 1963: geomagnetic polarity epochs and pleistocene geochronometry. nature 198, 1049– 1051. hald, n. & waagstein, r. 1984: lithology and chemistry of a 2km sequence of lower tertiary tholeiitic lavas drilled on suðuroy, faeroe islands (lopra-1). in: berthelsen, o., noe-nygaard, a. & rasmussen, j. (eds): the deep drilling project 1980–81 in the faeroe islands. annales societatis scientiarum faeroensis, supplementum ix, 15–38. tórshavn: føroya fróðskaparfelag. heirtzler, j.r., dickson, g.o., herron, e.m., pitman iii, w.c. & le pichon, x. 1968: marine magnetic anomalies, geomagnetic field reversals and motions of the ocean floor and continents. journal of geophysical research 73, 2119–2136. ito, h. & nogi, y. 1995: magnetic structures of seamounts in the western pacific ocean deduced from leg 144 downhole magnetometer logs. proceedings of the ocean drilling program, scientific results 144, 631–638. levanto, a.e. 1958: a three-component magnetometer for small drill-holes and its use in ore prospecting. geophysical prospecting 7, 183–195. parasnis, d.s. 1979: principles of applied geophysics, 275 pp. london: chapman & hall. pozzi, j.p., martin, j.p., pocachard, j., feinberg, h. & galdeano, a. 1988: in situ magnetostratigraphy: interpretation of magnetic logging in sediments. earth and planetary science letters 88, 357–373. pozzi, j.p., barthés, v., thibal, j., pocachard, j., lim, m., thomas, t. & pagès, g. 1993: downhole magnetostratigraphy in sediments: comparison with the paleomagnetism of a core. journal of geophysical research 98, 7939–7958. schönharting, g. & abrahamsen, n. 1984: magnetic investigations on cores from the lopra-1 drillhole, faeroe islands. in: berthelsen, o., noe-nygaard, a. & rasmussen, j. (eds): the deep drilling project 1980–81 in the faeroe islands. annales societatis scientiarum faeroensis, supplementum ix, 109–114. tórshavn: føroya fróðskaparfelag. vine, f.j. & matthews, d.h. 1963: magnetic anomalies over ocean ridges. nature 199, 947 only. waagstein, r. 1988: structure, composition and age of the faeroe basalt plateau. in: morton, a.c. & parson, l.m. (eds): early tertiary volcanism and the opening of the ne atlantic. geological society special publication (london) 39, 225–238. waagstein, r. & hald, n. 1984: structure and petrography of a 660 m lava sequence from the vestmanna-1 drill hole, lower and middle basalt series, faeroe islands. in: berthelsen, o., noenygaard, a. & rasmussen, j. (eds): the deep drilling project 1980–81 in the faeroe islands. annales societatis scientiarum faeroensis, supplementum ix, 39–70. tórshavn: føroya fróðskaparfelag. waagstein, r., guise, p. & rex, d. 2001: k/ar and 39ar/40ar wholerock dating of zeolite-facies metamorphosed flood basalts: the upper paleocene basalts of the faroe islands, ne atlantic. in: jolley, d.w. & bell, b.r. (eds): the north atlantic igneous province: stratigraphy, tectonic, volcanic and magmatic processes. geological society special publication (london) 197, 219–252. manuscript received 22 december 2000; revision accepted 13 september 2005. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1949 geological survey of denmark and greenland bulletin 6, 89-97 89geological survey of denmark and greenland bulletin 5, 89–97 © geus, 2004 jurassic dinoflagellate cysts from hochstetter forland, north-east greenland stefan piasecki and lars stemmerik three sections in hochstetter forland, north-east greenland, referred to the jurassic payer dal and bernbjerg formations, have been analysed for dinoflagellate cysts. the dinoflagellate cysts, new finds of ammonites and previously recorded marine faunas form the basis for improved dating of the succession. the basal strata of the payer dal formation at kulhus is here dated as late callovian, peltoceras athleta chronozone, based on the presence of relatively abundant limbicysta bjaerkei, mendicodinium groenlandicum, rhychoniopsis cladophora and tubotuberella dangeardii in an otherwise poor upper callovian dinoflagellate assemblage. ammonites have not been recorded from these strata. the upper payer dal formation at agnetesøelven is dated as late oxfordian, amoeboceras glosense – amoeboceras serratum chronozones, based on the presence of sciniodinium crystallinum, together with cribroperidinium granuligera and stephanelytron sp. the age is in accordance with ammonites present in the uppermost part of the formation at søndre muslingebjerg. new ammonites in the bernbjerg formation at agnetesøelven together with dinoflagellate cysts indicate an earliest kimmeridgian age, rasenia cymodoce and aulacostephanoides mutabilis chronozones. the upper callovian dinoflagellate cysts from hochstetter forland belong to a local brackish to marginal marine assemblage, which only allows a fairly broad correlation to coeval assemblages in central east greenland. in contrast, the oxfordian and kimmeridgian assemblages are fully marine and can be correlated from milne land in central east greenland via hochstetter forland to peary land in eastern north greenland. keywords: ammonites, boreal, dinoflagellate cysts, hochstetter forland, jurassic, north-east greenland geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: sp@geus.dk the northernmost onshore jurassic outcrops in east greenland occur in hochstetter forland and, some 60– 80 km farther north, on store koldeway (fig. 1); the stratigraphy of the latter area is described in an accompanying paper (piasecki et al. 2004a, this volume). the hochstetter forland peninsula is dominated by flat lowlands dissected by small streams; outcrops of pre-quaternary sediments are restricted to stream cuts and low coastal cliffs scattered throughout the region (fig. 1). however, at søndre muslingebjerg to the south, caledonian basement and middle–upper jurassic sandstones and coals are faulted and reach up to approximately 400 m above sea level. caledonian basement and jurassic sediments also crop out in northern hochstetter forland, east of agnetesø, and the meanders of agnetesøelven (the river from agnetesø to the coast) erode into marine jurassic sediments deposited on basement rocks (fig. 1). the jurassic stratigraphy of the hochstetter forland area is based on records of marine faunas including ammonites from several isolated exposures (ravn 1911; surlyk 1978). the lowermost deposits consist of the non-marine to marginal marine muslingebjerg formation, which is undated owing to the absence of marine fossils. the overlying sands of the payer dal formation contain upper oxfordian ammonites in the upper geus bulletin no 5.pmd 29-10-2004, 11:1489 90 part, and are followed by mudstones of the bernbjerg formation that yield upper oxfordian – lower kimmeridgian ammonites (surlyk 1978). the outcrops at kulhus and agnetesøelven expose the main part of the jurassic succession in hochstetter forland. these localities were visited during fieldwork in 1987 in order to collect material for palynological analysis with the aim of refining the stratigraphy. the present paper describes for the first time the dinoflagellate cyst floras from these northern outcrops close to the transition to the boreal dinoflagellate cyst province, for example in peary land (håkansson et al. 1981) and on svalbard (århus 1988). material at kulhus (fig. 1; locality 1), on the south coast of hochstetter forland west of søndre muslingebjerg, sand and coal seams were sampled during a ground stop on helicopter reconnaissance in 1987. only one sample (ggu 351570) contains dinoflagellate cysts. two closely situated localities at agnetesøelven were spotted from the air and visited during a short ground stop. the westernmost sandstone outcrop (locality 2) was measured and two fine-grained samples with at least some potential for palynology were collected from two horizons (fig. 2). eastwards and down-river, the sandstone-dominated succession was seen to be faulted against a succession of laminated mudstones (locality 3), and this was closely sampled for palynology (fig. 3). only a few, well-preserved dinoflagellate cysts were recovered from the sand succession at locality 2. by 20ºw 20 km 76º30'n undifferentiated, mostly glacial deposits cretaceous locality fault caledonian crystalline basementsøndre muslingebjerg kulhus nanok n hochstetter forland 1: milne land 2: jameson land 3: wollaston forland 4: hochstetter forland 5: store koldewey 6: germania land jurassic 7: peary land greenland 500 km agnetesø 2 3 7 1 2 3 4 56 1 1 agnetesøelven fig. 1. simplified sketch map of hochstetter forland illustrating the jurassic outcrops and the sampled localities. the map is modified from surlyk (1978, fig. 1). geus bulletin no 5.pmd 29-10-2004, 11:1490 91 contrast dinoflagellate cysts are abundant but poorly preserved in the mudstones at locality 3. ammonites are abundant in the lower, pyritic, part of the mudstone succession and in loose concretionary beds at the base of thecliff. these ammonites provide independent stratigraphical control of the dinoflagellate cyst assemblages. all of the samples from the three localities on hochstetter forland were prepared by traditional palynological methods and analysed for their content of dinoflagellate cysts. geology the jurassic succession on hochstetter forland is divided into the muslingebjerg, payer dal and bernbjerg formations of the vardekløft group. the marine sandstones of the payer dal formation were previously included in the pelion member of the vardekløft formation (sensu surlyk 1978); the pelion member is now raised to formation status (surlyk 2003, fig. 5). kulhus (locality 1) the low coastal cliff at kulhus, south-west hochstetter forland (fig. 1, locality 1), comprises 3–4 main coal belemnite bivalve oyster pecten spp. planolites ispp. wood bioturbation fine-grained sand planar lamination/bedding ggu sample fossils, etc. planar cross-lamination lenticular lamination trough cross-bedding scour and fill shell bed structures lithology hard bed mud sand 0 5 15 25 30 35 10 20 m 351538 351537 351537 mud pyrite ammonite fig. 2. sedimentological log of the upper payer dal formation at agnetesøelven (locality 2) showing the positions of analysed samples. geus bulletin no 5.pmd 29-10-2004, 11:1491 92 seams interbedded with black carbonaceous shales and light-coloured sandstones of the muslingebjerg formation (clemmensen & surlyk 1976). no marine fossils have been recorded from the formation but the high sulphur content of the coals and shales suggests a marine depositional environment (petersen et al. 1998). the jurassic succession at kulhus extends up the western flank of søndre muslingebjerg where the stratigraphically highest strata contain upper oxfordian ammonites referable to the amoeboceras glosense and/ or a. serratum zones (ravn 1911; sykes & surlyk 1976). coals and shales of the muslingebjerg formation havebeenprepared palynologically without finding any microscopic marine fossils. at kulhus, the first dinoflagellate cysts appear in the basal mudstone bed of the overlying payer dal formation, immediately above the highest coal seam, at the same level as the earliest marine faunas.restrictedassemblages of sporomorphs from the underlying coals indicate an overall jurassic age. agnetesøelven (localities 2 and 3) fine-grained sandstone, with abundant marine fossils, is exposed at locality 2 and referred to the payer dal formation (figs 1, 2). shell beds of pecten spp., oysters, other bivalves and serpulids occur throughout the succession and are commonly concentrated in scour fills. belemnites are present, but no ammonites were recovered during the short visit. woody material and small logs are also common. the sandstone is intensely bioturbated and many sedimentary structures are obliterated although cross-bedding or lamination is recognisable in most beds. planolites ispp. is common at certain horizons. the shale succession exposed at locality 3 is referred to the bernbjerg formation. it consists of laminated, dark mudstones alternating with lenticular-bedded, fine-grained, grey sandstones (fig. 3). ammonites and bivalves are abundant in a concretionary bed low in the section and belemnites and bivalves occur scattered higher in the succession. the concretionary beds are washed out from the lowermost succession and lie at the foot of the cliff. they contain accretions of ammonites in several stacked laminae together with abundant buchia sp. abundant male and female individuals occur together in the ammonite assemblages (j.h. 351546 351547 351548 351545 351544 351543 351542 351541 351540 351539 mud sand 0 5 10 15 20 25 m fig. 3. sedimentological log of the bernbjerg formation at agnetesøelven (locality 3) showing the positions of analysed samples. for legend, see fig. 2. geus bulletin no 5.pmd 29-10-2004, 11:1492 93 callomon, personal communication 1999). the assemblage is equivalent to ammonite fauna 15 from milne land in the lower rasenia cymodoce zone, lower kimmeridgian (fig. 4; birkelund & callomon 1985; j.h. callomon, personal communication 1999). lower kimmeridgian ammonites have been reported previously from sandstones at the locality of nanok in southern hochstetter forland (frebold 1932). ammonites of the lowermost kimmeridgian rasenia cymodoce and aulacostephanoides mutabilis zones, have been collected from mudstones of the bernbjerg formation in hochstetter forland (frebold 1932; surlyk 1978). stratigraphy basal payer dal formation (locality 1) dinoflagellate cysts. an unusual and relatively poor assemblage of dinoflagellate cysts was recorded immediately above the lithological transition from the barren muslingeelv formation to the fossiliferous payer dal formation (fig. 5). limbicysta bjaerkei, pilosidinium fensomei and pareodinia halosa dominate the assemblage, in association with gonyaulacysta jurassica, nannoceratopsis sp., occisucysta sp., pareodinia sp., solisphaeridium sp., tubotuberella cf. dangeardii and tubotuberella cf. egemenii. single specimens of atopodinium haromense, mendicodinium groenlandicum and rhynchodiniopsis cladophora were recorded. age. stratigraphically diagnostic species are few in this assemblage. in east greenland, limbicysta bjaerkei has not been recorded stratigraphically higher than the basal boreal middle jurassic in the cranocephalites borealis chronozone in jameson land (milner & piasecki 1996). however, a stratigraphical range of middle callovian and possibly into lowermost upper callovian has been reported from both the subboreal and arctic regions (smelror 1987, 1993). in jameson land, m. groenlandicum appears no lower than the kosmoceras jason chronozone (mid-callovian) and distinct rhynchodiniopsis cladophora and tubotuberella dangeardii appear in the basal upper callovian in the peltoceras athleta chronozone (milner & piasecki 1996). in conclusion, the sparse data indicate an age equivalent to the earliest late callovian, p. athleta chronozone (fig. 4). an early callovian age previously indicated for this unit (petersen et al. 1998) was based on a sample that was subsequently found to be from another section and locality. depositional environment. the dinoflagellate cyst assemblage is dominated by three species (fig. 5). limbicysta bjaerkei is possibly an acritarch because no clear archaeopyle has been documented. the assemblage differs markedly from normal marine assemblages described from time equivalent strata in milne land (piasecki 1996). bailey & hogg (1995) reported abundant l. bjaerkei in otherwise non-marine assemblages. this may indicate that the associated, frequent species, pilosidinium fensomei and pareodinia halosa, may have had similar environmental preferences. the abundance of l. bjaerkei – together with the restricted lithostratigraphy biostratigraphy formation ammonite zone age based on dinoflagellate cysts chronozones bernbjerg a. mutabilis r. cymodoce a. mutabilis r. cymodoce early kimmeridgian payer dal a. glosense – a. serratum a. glosense – a. serratumlate oxfordian late callovian p. athleta muslingebjerg no marine fossils (callovian?) pre-p. athleta chronostratigraphy fig. 4. summary of the jurassic stratigraphy of hochstetter forland. geus bulletin no 5.pmd 29-10-2004, 11:1493 94 assemblage – is taken as evidence for estuarine, brackish depositional environments during the initial transgression of hochstetter forland. upper payer dal formation (locality 2) dinoflagellate cysts. dinoflagellate cysts are relatively sparse in these sediments. however, the diversity is moderately good and the preservation is fine. ambonosphaera calloviense is the only dinoflagellate species representedbymore than one or two specimens in the assemblage (fig. 5). ambonosphaera calloviense, sirmiodinium grossii and sentusidinium sp. are the only species common to both samples. age. despite the paucity of dinoflagellate cysts, the co-occurrence of sciniodinium crystallinum, cribroperidinium granuligera and stephanelytron sp. indicates a late oxfordian age i.e. amoeboceras glosense to amoeboceras serratum chronozones, by comparison to hold with hope and milne land further to the south (piasecki 1996; piasecki et al. 2004a, b, this volume; figs 1, 4). none of the other dinoflagellate cysts are inconsistent with this age, which is also in accordance with the age indicated by ammonites from the uppermost payer dal formation at søndre muslingebjerg (ravn 1911; sykes & surlyk 1976). however, the dinoflagellate cyst assemblage is too restricted to allow a more precise correlation. depositional environment. the low abundance combined with the moderate diversity of dinoflagellate cysts indicates near-shore marine deposition in a high-energy environment. the sandy sediments with a rich benthic fauna, partly in situ and partly reworked into shell beds, support this interpretation. bernbjerg formation (locality 3) dinoflagellate cysts. dinoflagellate cysts are abundant and relatively diverse, but their preservation is poor. the composition of the assemblages varies significantly through the relatively short section (fig. 5). extremely abundant sirmiodinium grossii characterises the lower part of the section and is gradually replaced by abundant gonyaulacysta jurassica sensu lato in the upper part. five other species appear in succession with distinct and characteristic maxima through the succession (fig. 5). abundant nummus sp. occurs together with the maximum numbers of escharisphaeridium pocockii in the lower part of the succession. this is followed closely by a maximum abundance of cribroperidinium granuligera, then by a maximum of perisseiasphaeridium pannosum (together with gonyaulacysta jurassica sensu lato) and finally by maximum abundance of occisucysta cf. monoheuriskos in the uppermost sample. a comparable succession of dinoflagellate cyst assemblages has been recorded across the boundary of the payer dal and bernbjerg formations at kløft ii on store koldewey (piasecki et al. 2004a, this volume). at this locality, a maximum of g. jurassica sensu lato is followed by abundant p. pannosum and occisucysta cf. monoheuriskos, similar to that recorded from the upper part of the section at agnetesøelven (locality 3). assemblages dominated by nummus sp., e. pocockii and c. granuligera were not recorded at store koldewey, but these species are present. in contrast, the stratigraphically significant paragonyaulacysta capillosa is abundant at store koldewey, whereas it is rare in the succession at agnetesøelven. age. the succession cannot be older than early kimmeridgian based on the ammonite fauna in the basal strata,which is indicative of the lower rasenia cymodoce zone. this is in accordance with the dinoflagellate cyst assemblage of abundant sirmiodinium grossii and frequent gonyaulacysta jurassica, adnatosphaeridium sp. and paragonyaulacysta capillosa. to the south, on milne land, p. capillosa first appears in the r. cymodoce chronozone and is followed by perisseiasphaeridium pannosum in the succeeding aulacostephanoides mutabilis chronozone (piasecki 1996). at agnetesøelven, the appearance of p. pannosum higher in the succession accordingly indicates an age equivalent to the a. mutabilis chronozone for this part of the succession (fig. 4). dinoflagellate cysts indicative of younger jurassic strata were not recorded. lower kimmeridgian ammonites have previously been recorded from the bernbjerg formation on hochstetter forland. an ammonite fauna referable to the a. mutabilis zone has been recovered from the nanok and agnetesøelven regions (frebold 1932; surlyk 1978). depositional environment. the dinoflagellate cysts are strongly degraded. they are physically broken, and angular imprints of crystals and deep circular imprints of spherical pyrite framboids obscure the sculpture and structure of their walls. together with the undisturbed lamination of the sediments, this indicates deposition geus bulletin no 5.pmd 29-10-2004, 11:1494 95 50250 66.00 61.00 56.00 53.00 50.00 47.00 45.00 40.00 25.00 20.00 1.00 351548 351547 351545 351544 351543 351542 351541 351539 351538 351537 351570 jurassic oxfordian–kimmeridgiancallovian bernbjergpayer dal 1 veryhachium spp. 2 tubotuberella dangeardii 3 atopodinium haromense 4 sentusidinium rioultii 5 mendicodinium groenlandicum 6 dissiliodinium spp. 7 botryococcus spp. 8 limbicysta bjaerkei 9 pilosidinium fensomei 10 solisphaeridium spp. 11 tubotuberella cf. egemenii 12 occisucysta spp. 13 nannoceratopsis spp. 14 pareodinia cf. stegasta 15 valensiella spp. 16 rhynchodiniopsis cladophora 17 pareodinia spp. 18 gonyaulacysta jurassica 19 pareodinia halosa 20 stephanelytron spp. 21 tenua spp. 22 ambonosphaera calloviense 23 sentusidinium spp. 24 sirmiodinium grossii 25 scriniodinium crystallinium 26 cribroperidinium granuligera 27 pareodinia cf. pachyceras 28 scriniodinium spp. 29 tubotuberella apatela 30 avellodinium cf. falsificum 31 rhynchodiniopsis cf. cladophora 32 scriniodinium cf. crystallinium 33 leptodinium spp. 34 leptodinium subtile 35 cribroperidinium sp. 36 paragonyaulacysta cf. capillosa 37 escharisphaeridia pocockii 38 nummus spp. 39 atopodinium spp. 40 occisucysta cf. monoheuriskos 41 adnatosphaeridium spp. 42 pareodinia stegasta 43 epiplosphaera spp. 44 sentusidinium pelionense 45 barbatacysta spp. 46 perisseiasphaeridium pannosum 47 paragonyaulacysta capillosa 48 paragonyaulacysta spp. 49 prolixosphaeridium granulosum 50 tubotuberella rhombiformis a lph a bet ic a l spec ies list 41 adnatosphaeridium spp. 22 am bonosphaera calloviense 3 atopodinium harom ense 39 atopodinium spp. 30 avellodinium cf. falsificum 45 barbatacysta spp. 7 botryococcus spp. 19 pareodinia halosa 26 cribroperidinium granuligera 35 cribroperidinium sp. 6 d issiliodinium spp. 43 epiplosphaera spp. 37 escharisphaeridia pocockii 18 g onyaulacysta jurassica 33 leptodinium spp. 34 leptodinium subtile 5 m endicodinium groenlandicum 13 n annoceratopsis spp. 38 n um m us spp. 12 o ccisucysta spp. 40 o ccisucysta cf. m onoheuriskos 48 paragonyaulacysta spp. 47 paragonyaulacysta capillosa 36 paragonyaulacysta cf. capillosa 27 pareodinia cf. pachyceras 14 pareodinia cf. stegasta 17 pareodinia spp. 42 pareodinia stegasta 8 lim bicysta bjaerkei 46 perisseiasphaeridium pannosum 9 pilosidinium fensom ei 49 prolixosphaeridium granulosum 31 rhynchodiniopsis cf. cladophora 16 rhynchodiniopsis cladophora 32 scriniodinium cf. crystallinium 25 scriniodinium crystallinium 28 scriniodinium spp. 44 sentusidinium pelionense 4 sentusidinium rioultii 23 sentusidinium spp. 24 sirm iodinium grossii 10 solisphaeridium spp. 20 stephanelytron spp. 21 tenua spp. 29 tubotuberella apatela 11 tubotuberella cf. egem enii 2 tubotuberella dangeardii 50 tubotuberella rhom biform is 15 valensiella spp. 1 veryhachium spp. sample height metre ggu sample no. system stage formation h ochstetter forland k ulhus and a gnetesøelven > 50 specim ens 20–50 specim ens 5–19 specim ens 1–4 specim ens fig. 5. stratigraphical distribution chart of dinoflagellate cysts in samples from all three localities. the sample heights indicated are largely arbitrary, although the three localities are arranged in stratigraphic order. sample at 1 m is from locality 1, samples at 20 m and 25 m are from locality 2 (arbitrary spacing, see fig. 2 for correct locations) and samples at 40–66 m are from locality 3 (spacings approximately to scale, see fig. 3 for precise locations). the recorded species are arranged by their first stratigraphical appearance. geus bulletin no 5.pmd 29-10-2004, 11:1495 96 below wave base in a low-oxygenated environment. lenticular laminae with small-scale ripples, however, show that the bottom water was not completely stagnant and that the sea floor was periodically swept by weak bottom currents. correlation studies of the jurassic ammonite and dinoflagellate cyst stratigraphy in east greenland, combined with sedimentological studies and sequence stratigraphical interpretations, contribute towards an integrated model in which units can be identified by their content of dinoflagellate cysts. the present study of the dinoflagellate cyst assemblages on hochstetter forland contributes basic data to this complex study. the assemblage dominated by limbicysta bjaerkei in the basal payer dal formation at kulhus has not been recorded anywhere else in east greenland. this assemblage comprises the first marine fossils to have been deposited above the coal-bearing floodplain environment of the muslingebjerg formation. the dinoflagellate cyst assemblage records deposition in a marginal marine to brackish environment. the overlying sandstones are interpreted as tidal facies followed by shoreface facies (petersen et al. 1998), reflecting a rise in relative sea level. this assemblage thus characterises marginal marine environments at the feather-edge of the jurassic depositional basin during a major drowning event (alsgaard et al. 2003). the poor dinoflagellate cyst assemblage in the upper payer dal formation is not representative of this stratigraphic interval, compared to the assemblages recorded from other localities in east greenland. however, it is associated with a stratigraphic unit equivalent to the a. glosense and a. serratum chronozones that previously has been identified throughout east greenland, i.e. in milne land, jameson land and hold with hope (engkilde 1994; piasecki 1996; vosgerau et al. 2004, this volume). the dinoflagellate cyst assemblage from the bernbjerg formation at agnetesøelven is known from milne land in the south to store koldewey in the north, and a similar assemblage occurs in the ladegårdsåen formation in peary land, north greenland (fig. 1; piasecki 1966; piasecki et al. 2004a, this volume). the dinoflagellate cyst assemblage correlates with the r. cymodoce and a. mutabilis chronozones. it is associated with a major kimmeridgian flooding event which allowed more permanent shelf anoxia to spread to shallow shelf areas of east greenland (milne land, wollaston forland, hold with hope, store koldewey) and north greenland (peary land). maximum flooding occurred in the a. mutabilis to a. eudoxus chrons. conclusions new ammonite data confirm and refine earlier age determinations of the jurassic succession on hochstetter forland. the age of the upper payer dal formation in hochstetter forland is confirmed as late oxfordian, and the age of the bernbjerg formation is confirmed as earliest kimmeridgian. the dinoflagellate cyst stratigraphy from the three localities is fragmentary. the assemblage from the lower payer dal formation at kulhus has not been reported from any other section in east greenland; it is important because it represents marginal marine conditions associated with a major flooding event of earliest late callovian age, p. athleta chronozone (fig. 4). this limits the age of the muslingeelv formation upwards. the assemblage in the upper payer dal formation is restricted, but supports the previously recorded late oxfordian age, corresponding to the a. glosense and/ or a. serratum chronozones. the third assemblage is well known from east greenland and dates the bernbjerg formation at agnetesøelven to the earliest kimmeridgian, r. cymodoce and a. mutabilis chronozones (fig. 4). this assemblage is associated with a major transgressive event characterised by extensive shelf anoxia in east greenland. acknowledgements the work was initiated as part of the project ‘resources of the sedimentary basins of north and east greenland’ supported by the danish research councils and completed by support from the carlsberg foundation ans. 980089/0-262. john h. callomon is thanked for the identification of the ammonites. the referees, d.j. batten and j.b. riding, provided constructive and very helpful suggestions. geus bulletin no 5.pmd 29-10-2004, 11:1496 97 references alsgaard, p.c., felt, v.l., vosgerau, h. & surlyk, f. 2003: the jurassic of kuhn ø, north-east greenland. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 865–892. århus, n. 1988: palynostratigraphy of some bathonian–hauterivian sections in the arctic, with emphasis on the janusfjellet formation type section, spitsbergen. institutt for kontinentalsokkelundersøkelser (iku) report 23.1252.11/01/88, 139 pp. bailey, d.a. & hogg, n.m. 1995: fentonia bjaerkei gen. et comb. nov.; transfer from parvocysta bjaerke 1980. journal of micropalaeontology 14, 58 pp. birkelund, t. & callomon, j.h. 1985: the kimmeridgian ammonite faunas of milne land, central east greenland. bulletin grønlands geologiske undersøgelse 153, 56 pp. clemmensen, l.b. & surlyk, f. 1976: upper jurassic coal-bearing shoreline deposits, hochstetter forland, east greenland. sedimentary geology 15, 193–211. engkilde, m. 1994: the middle jurassic vardekløft formation, east greenland: depositional environments and sequence stratigraphy of shallow marine sandstones deposited in a low-gradient epeiric seaway, 1–207. unpublished ph.d. thesis, university of copenhagen, denmark. frebold, h. 1932: geologie der jurakohlen des nördlichen ostgrönland. meddelelser om grønland 84(5), 62 pp. håkansson, e., birkelund, t., piasecki, s. & zakharov, v. 1981: jurassic–cretaceous of the extreme arctic (peary land, north greenland). bulletin of the geological society of denmark 30, 11–42. milner, p.s. & piasecki, s. 1996: boreal middle jurassic dinoflagellate cyst stratigraphy of jameson land, east greenland. in: piasecki, s. et al. (eds): formation of source and reservoir rocks in a sequence stratigraphic framework, jameson land, east greenland. energy research programme efp-93, projects 1313/93-0010 and 0017. danmarks og grønlands geologiske undersøgelse rapport 1996/30, vol. i and ii, 46 pp. petersen, h.i., bojesen-koefoed, j., nytoft, h.p., surlyk, f., therkelsen, j. & vosgerau, h. 1998: relative sea-level changes recorded by paralic liptinite-enriched coal facies cycles, middle jurassic muslingebjerg formation, hochstetter forland, northeast greenland. international journal of coal geology 36, 1–30. piasecki, s. 1996: boreal dinoflagellate cyst stratigraphy of middle to upper jurassic sediments of milne land, east greenland. in: piasecki, s. et al. (eds): formation of source and reservoir rocks in a sequence stratigraphic framework, jameson land, east greenland. energy research programme efp-93, projects 1313/93-0010 and 0017. danmarks og grønlands geologiske undersøgelse rapport 1996/30, vol. i and ii, 100 pp. piasecki, s., callomon, j.h. & stemmerik, l. 2004a: jurassic dinoflagellate cysts from store koldewey, north-east greenland. in: stemmerik, l. & stouge, s. (eds): the jurassic of northeast greenland. geological survey of denmark and greenland bulletin 5, 99–112 (this volume). piasecki, s., larsen, m., therkelsen, j. & vosgerau, h. 2004b: jurassic dinoflagellate cyst stratigraphy of hold with hope, north-east greenland. in: stemmerik, l. & stouge, s. (eds): the jurassic of north-east greenland. geological survey of denmark and greenland bulletin 5, 73–88 (this volume). ravn, j.p.j. 1911: on jurassic and cretaceous fossils from northeast greenland. meddelelser om grønland 45(10), 437–500. smelror, m. 1987: bathonian and callovian (middle jurassic) dinoflagellate cysts and acritarchs from franz josef land, arctic soviet. polar research 5, 221–238. smelror, m. 1993: biogeography of bathonian to oxfordian (jurassic) dinoflagellates: arctic, nw europe and circum-mediterranean regions. palaeogeography, palaeoclimatology, palaeoecology 102, 121–160. surlyk, f. 1978: mesozoic geology and palaeogeography of hochstetter forland, east greenland. bulletin of the geological society of denmark 27, 73–87. surlyk, f. 2003: the jurassic of east greenland: a sedimentary record of thermal subsidence, onset and culmination of rifting. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 659–722. sykes, r.m. & surlyk, f. 1976: a revised ammonite zonation of the boreal oxfordian and its application in north-east greenland. lethaia 9, 421–436. vosgerau, h., larsen, m., piasecki, s. & therkelsen, j. 2004: a new middle–upper jurassic succession on hold with hope, north-east greenland. in: stemmerik, l. & stouge, s. (eds): the jurassic of north-east greenland. geological survey of denmark and greenland bulletin 5, 51–71 (this volume). geus bulletin no 5.pmd 29-10-2004, 11:1497 wire-line log-based stratigraphy of flood basalts from the lopra-1/1a well, faroe islands 7 wire-line log-based stratigraphy of flood basalts from the lopra-1/1a well, faroe islands lars o. boldreel the present study shows that it is possible to use conventional borehole logs to perform a detailed lithological/stratigraphical division of a column of subaerially extruded basalt. a stratigraphical division of the subaerial flood basalts penetrated by the lopra-1/1a well has been carried out using new wire-line logging data measured in 1996 in the interval 200–2489 m depth. resistivity data acquired in the interval 200–2178 m depth during 1981 after the initial drilling of the lopra-1 well have also been incorporated. eighty-six individual flow units, 18 compound flows and two dolerite dykes have been identified by combining the nphi porosity, rhob density, p-, sand stonely-sonic transit time, calliper and resistivity logs. fifty-two sedimentary/tuffaceous layers have also been identified using the cgr and sgr gamma ray and potassium logs in combination with the aforementioned logs. within the flow units, sonic velocity, density and resistivity are highest in the core where porosity is lowest. this relation is reversed in the uppermost and basal zones of the flow units. the sonic velocity in the core seems to be independent of the thickness of the flow unit. porous zones seem abundant in some cores and the total section of cores containing porous zones constitutes more than 70% of the thickness of its flow unit, but where porous zones are absent the core makes up only roughly 50% of the thickness of the flow. it is suggested that the flow units with porous cores represent aa flows (88% of the flow units) and the others pahoehoe flows (12% of the flow units). the log pattern of the flow units (crust, core and basal zone) is similar to log patterns reported from other basalt plateaux. however the patterns in lopra-1/1a show a larger variation than elsewhere, suggesting that the flow units are more complex vertically than previously thought. statistical analysis of p-, sand stonely-waves, rhob, nphi, resistivity, gamma and calliper logs has been carried out. cross-plots based on the lithological divisions have been produced that show a pronounced reduction in scattering versus p-sonic transit time and pand s-sonic, rhob and nphi logs correlate with depth. the geochemical logs do not reflect the cyclic structure of the flow units and probably represent the primary composition of the basalt. the thorium log especially indicates flow units with high and low radioactivity and it is suggested that a minimum of 36 flow fields form the logged part of the lower basalt series. dolerite units described in previous works have been confirmed based on the combined interpretation of wire-line logs. the log data suggest that the subaerially extruded basalt has its base at a depth of approximately 2490 m and that a hyaloclastite succession is found below that depth. the transition from subaerially extruded basalt to hyaloclastite produces a negative acoustic impedance and it is found that the transition corresponds to a negative reflection interpreted on vsp surveys from 1988 and 1994. keywords: flood basalts, flood basalt petrophysics, north atlantic, faroe islands, lopra-1/1a well, faroese basalt lithology, wire-line logs _________________________________________________________________________________________________________________ geological institute, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: lob@geol.ku.dk © geus, 2006. geological survey of denmark and greenland bulletin 9, 7–22. available at: www.geus.dk/publications/bull geus bulletin no 9 7 juli.pmd 07-07-2006, 14:197 8 the lopra-1 well was deepened in 1996 to drill through a pronounced negative seismic reflection at an estimated depth of 2366 m and an additional reflection at a depth of 3486 m found from the analysis of the data acquired in vertical seismic profiling (vsp) surveys carried out in 1988 (kiørboe & petersen 1995) and 1994. several wireline logs were run in the upper part of the well (0–2184 m) that had been drilled in 1981 to supplement the previous logging (nielsen et al. 1984). a more extensive logging programme was carried out in the deepened part of lopra1/1a and the new wire-line data for the entire well have been analysed. the wire-line logs were used to investigate the lower basalt series of the faroe islands, to construct a lithological and stratigraphical division of the lava sequence and to test whether a subdivision of the basalt column would result in better statistics of its physical properties. the results of the new logging have been both new understanding and results in greater detail than was possible from the data available from the 1981 drilling and logging, and they supplement studies carried out in other basalt covered regions in the world. it was intended that analysis of the wire-line logs acquired in 1996 should include the logs acquired in 1981, but this turned out to be complicated. it was possible to use the older resistivity data only after they had been corrected manually within individual flow units. consequently this paper is concerned mainly with interpretation of the wire-line log data measured in 1996 in the subaerially extruded lower basalt series. geology of the faroe islands the faroe islands consist of subaerially extruded tholeiitic basalt flows erupted in the late paleocene in connection with the opening of the northern north atlantic ocean. the basalts consist of three lava formations called the upper, middle and lower basalt formations (or series). a tuff-agglomerate zone and a coal-bearing horizon separate the lower and the middle formations (e.g. rasmussen & noe-nygaard 1969, 1970; 1990). the three lava formations are all exposed on the islands and the lowest exposed stratigraphical level is found in the southern part of the southernmost island, suðuroy, where the original lopra-1 well was drilled. geological and geophysical investigations based on data from the lopra-1 well resulted in six papers published in berthelsen et al. (1984). the geology of the well was described by hald & waagstein (1984) who found the lava to consist of tholeiitic basalt. a stratigraphical division of the well was established on the basis of drill cuttings collected every 2 m, five cores, the wireline logs (acquired in 1981) and geochemical rock analysis. this resulted in a division of the sequence into 87 flow units, 27 sedimentary units and two dolerite dykes within the interval 200–2180 m depth plus an additional 10 flow units and five sedimentary beds in the cased interval 0–200 m. the bedding in the well was established using chilled vesicular (amygdaloidal) basalt as a marker of flow tops. the observed transition between vesicular and non-vesicular basalt was afterwards correlated to the neutron-neutron and resistivity logs (waagstein et al. 1982; nielsen et al. 1984). the flow units range in thickness from a few metres to about 50 m with the majority of flows having thicknesses of 5–10 m or 10–15 m. most of the flows were interpreted as aa type based on the characteristics of the topmost cuttings. sediments were observed on the top of about one third of the drilled lava flows. two dolerite dykes were found at 834–738 m and 616– 508 m, and hald & waagstein (1984) suggested that the real thickness of both dykes is about 9 m assuming an inclination of 85°. the volcaniclastic sediments were distributed more or less randomly throughout the drilled sequence, and no correlation was found either with flow thickness or with chemical composition of the flows. it was concluded that there were no major hiatuses during eruption of the drilled sequence or the overlying 700 m of exposed lower basalt formation. data from the lopra-1 well the original lopra-1 stratigraphic well reached a total depth (td) of 2184 m in 1981. it was deepened in 1996 table 1. wire-line log suites 1, 2 and 5 used in this study suite 1: 196–2184 m (in the existing hole drilled in 1981). two log runs were carried out: run no. 1: ldl-cnl-ngs (recording a maximum hole deviation of 1º from the vertical). gr was run to the surface run no. 2 pass 1: dsi-cross and u&l dipole run no. 2 pass 2: dsi-p&s and stoneley. suite 2: 2185–3158 m three runs were carried out: run no. 1: ldl-cnl-ngs-dsi. run no. 2: shdt-gr-ams recording a maximum hole deviation from the vertical of 14° run no. 3: ghmt-gr. suite 5: 2204–3531 m run no 1: msct-gr geus bulletin no 9 7 juli.pmd 07-07-2006, 14:198 9 to a td of 3158 m and was sidetracked from 3091 m after a partly unsuccessful fishing operation. the sidetrack (lopra-1a) well reached td at 3565 m and the lopra-1/1a well sections were plugged and abandoned after final wire-line logging. the lopra-1/1a well has been logged on various occasions. caliper, gamma ray, resistivity, neutron-neutron porosity and sonic logs were acquired in 1981 (nielsen et al. 1984), but the sonic log and part of the resistivity log turned out to be unsuccessful (nielsen et al. 1984). successful vsp surveys were carried out in 1988 (kiørboe & pedersen 1995) and 1994, and ahead-of-tool predictions were carried out in order to evaluate the remaining thickness of the basalt. the deepened well was logged in 1996 and a large number of logs, including a vsp, were obtained. a total of five suites were run between 196 and 3565 m depth. suites 1 to 5 were acquired in the intervals 196–2184 m, 2185–3158 m, 3091–3516 m, 3516–3561 m and 2204–3531 m, respectively. three of the suites have been used in this study, suites 1, 2 and 5 (table 1). from the suite of wire-line logs, the calliper, sgr (total gamma ray), cgr gamma ray (computed gamma ray, including measurements of thorium and potassium), p-, sand stonely-wave transit time, rhob (bulk density), nphi (neutron porosity hydrogen index), resistivity (short and long from 1981), potassium, thorium and uranium logs were selected for this study. after processing (apart from the resistivity logs) by schlumberger, the data were interpreted on an sgi unix workstation using landmark software (stratworks® and petroworks®). there were problems with the logs acquired by the spectral gamma log tool in the deepened part of the lopra-1 well, probably because calibration had not been carried out properly. the data were corrected by shifting the thorium, potassium and cgr logs to the responses measured in the previously drilled hole. an attempt was made to correlate the wireline logs from 1981 and 1996, as both suites include a calliper. that task proved not to be simple, probably because of different cable tensions and depth annotations. it is felt that the 1981 resistivity data have been correctly adjusted by manual editing and adjusting the 1981 data to the interpretation of the wire-line suite obtained in 1996. characteristics of subaerially extruded basalt it is known from various workers (e.g. cashman & kauahikaua 1997; self et al. 1997; thordarson & self 1998; waagstein 1999) that subaerially extruded lava flows have some characteristics in common and the following discussion focuses on those characteristics which may be expected to be confirmed from the logging suite run in the lopra-1 well. a lava formed in a single volcanic eruption is called a flow field and consists of one or several lava flows erupted more or less continuously from the same vent area or fissure. the commonest diagnostic criterion likely to be used is the chemical composition of the basalt magma that is assumed to be fairly constant during an eruption. a lava flow may consist of several flow units, partly or completely surrounded by chilled crust. a flow unit consists of a lava crust, a lava core and a basal zone. the appearance of the crust has been used to define two morphological types, aa and pahoehoe, in both of which the crust is usually much thicker than the basal zone. the core typically consists of massive basalt with few or no vesicles, and sometimes displays flow banding. thick cores often show a more or less columnar jointing in the lower part which propagated upwards perpendicular to the base of the flow during cooling. thinner and less regular columns in the upper part may have propagated downwards from the top of the flow. vesicular zones in the core have been reported by mcmillan et al. (1987) who showed a two-stage vesiculation of basalt from the colombia river basalt plateau. walker (1989) studied vesicle distribution in profiles across flow units, on hawaii primarily of pahoehoe type but also of aa type, and concluded that the distribution of vesicles does not need to be symmetrical about a horizontal median plane. the rubble near the surface of a flow unit in aa lava may be unconsolidated with fractures and voids filled by tuff, and sometimes the rubble fragments are welded. the fragmentation decreases downwards, and a complete gradation into unbrecciated basalt in the core of the flow unit is commonly seen. thick flow units in the lower formation on the faroes are typically from several metres to several tens of metres thick, as seen on suðuroy (waagstein 1999). thick flow units usually have an upper rubbly crust several metres thick whereas the lower crust or basal zone is generally less than one metre. the vesicles and fractures are usually filled with secondary minerals. aa lavas on hawaii normally consist of a large number of thin flow units, 0.2–2.0 m thick. pahoehoe type lava has a continuous vesicular crust that is distinguished from its core mainly by its vesicularity, and roughly half of the total thickness of a pahoehoe flow consists of crust. the crust is highly vesicular near its top surface with abundant small vesicles that often exceed 50% of the bulk rock volume. the vesicles show an overall decrease in abundance and increase in size downwards reflecting the increase in gas pressure exerted by the thickening lava crust. the vesicles are often arranged in horigeus bulletin no 9 7 juli.pmd 07-07-2006, 14:199 10 zontal layers with a vertical spacing of about 10 cm, especially in the upper part of the crust. horizontal vesicular veins or sheets of basalt up to about 10 cm thick may occur near the top of the core, but otherwise the core of the flow is usually massive with no or only a few relatively large vesicles. the basal vesicular zone is usually a few tens of centimetres thick at most and sometimes sub-vertical pipes are found within it. sediments thin beds of tuff with a vertical extent of between a few centimetres and few tens of metres are found within the faroes lava succession. the beds are typically very finegrained, clayey, fissile with vivid colours. the tuff beds are buried soils (palaeosols) formed by disintegration and chemical break down of volcanic materials. the beds may originate from either airfall volcanic ash or from in situ weathering of the lava flows (hald & waagstein 1984; waagstein 1999). microfossils suggest that the latter formed slowly enough to allow immigration of plants and other living organisms before being covered by the next lava flow, thus representing quiescence in the erupting environment (lund 1983). previous work on the relationship between lithology and logging response the geophysical response of interbedded sediment/tuff layers and subaerially extruded basalt, characterised both by the rhythmic succession of flow units and the subdivision of the flow unit into three parts, may be examined from wire-line logging data and drilling penetration rate. this has been done by a number of authors from various places around the world: the faroe islands (kern & richter 1979; nielsen et al. 1984), odp hole 642 (planke & flovenz 1996), odp hole 642e (planke 1994), newark basin (goldberg et al. 1994), north sea basalt flows from the middle jurassic (rider 1996), deccan trap area (buckley & oliver 1990; singh 1996), japan and chile (berlitz et al. 1988) and a geothermal well with no cited geographical location (sanyal et al. 1980). physical parameters obtained from the analyses of wire-line data by these workers are presented in table 2, and the results from the present study are found in table 3. in two studies where the layered basalt flows represent the seaward dipping reflector sequence, from odp hole 917 (which also had formation microscanner images) (planke & flovenz 1996) and from odp hole 642e (planke 1994), the interpretation of wire-line logs was integrated with continuous coring. the above studies suggest that the following features are of special interest in relation to the wire-line logs obtained from lopra-1/1a. the often vesicular, fractured and altered crust and flow tops are characterised by gradients in the values represented * chile; † japan; # uncorrected values; § average cps: counts per second; p.u.: neutron porosity units; gapi: gamma log in api units; mas: the massive core of a flow unit; por: the porous crust of a flow unit; sed: sediments between flow units. kern & richter (1979) faroese basalt berlitz et al. (1988) chile japan goldberg et al. (1994) usa planke (1994) odp hole 642e planke & flovenz (1996) odp hole 917 singh (1996) deccan traps 10 p.u. 3.06*; 2.8† 15 3–5.5 4§ 2.5§ 6–8 m 4.79–6.21 2.73–3.52 2.73–3.52 christensen (1996) 5.91§ 3.03§ 2.882§ 6.2§ 10%# 3.05§ 75 ± 25 5.0–6.0 3.0–5.3 < 5% corrected 15–30% corrected 10–20% corrected 2.3–2.8 5 gapi 20 gapi 15–40 cps 40–100 cps 100–300 cps table 2. physical properties of basalt based on wire-line logs vp (km/sec) vp massive (km/sec) vp sediments (km/sec) vs (km/sec) porosity porosity massive porosity porous porosity sediment density (g/cm3) gamma-ray (gapi) gamma-ray (mas) gamma-ray (por) gamma-ray (sed) unit thickness 6–8 m geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1910 11 vp flow unit vp core vp crust vp compound flow vp dolerite dyke vp sediments vs flow unit vs core vs crust vs compound flow vs dolerite dyke vs sediments vst flow unit vst core vst crust vst compound flow vst dolerite dyke vst sediments porosity flow unit porosity core porosity crust porosity compound flow porosity dolerite dyke porosity sediments density flow unit density core density crust density compound flow density dolerite dyke density sediments gamma ray flow unit gamma ray core gamma ray crust gamma ray compound flow gamma ray dolerite dyke gamma ray sediments calliper flow unit calliper core calliper crust calliper compound flow calliper dolerite dyke calliper sediments resistivity (lld) flow unit resistivity (lld) core resistivity (lld) crust resistivity (lld) compound flow resistivity (lld) dolerite dyke resistivity (lld) sediments resistivity (ild) flow unit resistivity (lls) core resistivity (lls) crust resistivity (lls) compound flow resistivity (lls) dolerite dyke resistivity (lls) sediments 5.53 5.83 4.69 4.92 6.08 4.26 2.97 3.14 2.52 2.59 3.28 2.34 1.46 1.47 1.42 1.44 1.45 1.40 0.19 0.15 0.31 0.30 0.12 0.31 2.85 2.90 2.66 2.74 2.90 2.58 10.40 10.23 9.75 10.32 4.74 27.09 10.47 10.52 10.15 10.26 11.81 11.84 11.96 11.96 11.73 11.80 12.12 12.08 average table 3. physical properties of basalt based on wire-line logs from lopra-1 317.16 369.51 165.68 167.66 515.80 91.91 558.39 661.47 283.45 (37.35–1777.40) 1038.07 288.27 356.15 122.45 121.70 * limestone porosity units; om: ohms m2/m 1055.64 601.04 187.91 144.09 9.50 (0.7–61.1) api 9.39 (0.0–33.77) 9.43 (0.0–33.76) 9.18 (1.34–36.53) 4.57 (1.36–26.95) (1.66–78.28)25.21 2.96 2.93 2.66 2.76 2.95 2.64 0.32 0.10 0.30 0.32 0.14 0.16 1.41 1.47 1.44 1.43 1.48 1.47 2.34 3.39 2.60 (1.90–3.34) (1.90–4.03) (1.82–3.44) 2.54 (1.74–3.82) 3.19 3.11 4.30 6.29 5.02 4.73 5.90 5.78 473.36 527.20 76.63 (24.07–1777.40) om (6.64–15.83) inches (6.64–15.83) (8.24–15.62) (10.10–14.61) (9.49–18.58) (10.09–18.41) (2.03–3.01) (1.51–3.04) (2.53–2.95) (1.55–3.15) (1.57–3.15) (1.55–3.15) g/cm3 (0.11–0.50) (0.04–0.45) (0.15–0.49) (0.15–0.44) (0.04–0.47) (0.04–0.50) % lspu* (1.19–1.49) (1.28–1.48) (1.19–1.49) (1.11–1.50) (1.22–1.52) (1.11–1.52) km/sec (2.00–4.01) (1.74–4.01) km/sec (2.86–6.31) (3.37–7.58) (2.43–6.57) (2.94–6.63) (2.94–7.47) (2.94–7.47) km/sec median (min–max) 164.21255.35 (22.37–309.21) (22.37–760.66) (26.32–596.06) (18.52–731.21) (32.43–766.68) (18.52–766.66) om (26.74–890.36) (145.01–1773.44) (52.07–1135.72) (24.67–1197.98) 178.05 geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1911 12 by the porosity dependent wire-line logs. the p-wave velocity, density and resistivity values are rather low but increase into the flow, in contrast to the porosity log that shows rather high values that decrease into the flow (planke & flovenz 1996). natural gamma counts are intermediate (goldberg et al. 1994; planke 1994; rider 1996; singh 1996). the crust is often soft and weathered and can be recognised from an increase in the rate of drilling pentration (buckley & oliver 1990). the massive core is characterised by rather high values of p-wave velocity, density and resistivity in contrast to rather low values of porosity and low natural gamma counts (goldberg et al. 1994; planke 1994; planke & flovenz 1996; rider 1996; singh 1996). a number of faroes basalt samples, selected to be free of fractures and secondary alteration, showed in laboratory experiments that the pand s-wave velocities increase as a function of pressure (kern & richter 1979). two distinct gamma ray count rates of primary potassium and thorium show that high-gamma basalts exist (buckley & oliver 1990). the basal zone is characterised by gradients in values from the wire-line logs representing porosity. p-wave velocity, density and resistivity decrease rapidly downwards whereas porosity increases downwards near the base of the unit (goldberg et al. 1994; planke 1994; planke & flovenz 1996). the sediments interbedding the flow units show very low resistivity and very high gamma counts (nielsen et al. 1984; singh 1996) and intermediate porosity and density values (planke 1994). nielsen et al. (1984) showed that different series within a large vertical section of basalt flows could be identified primarily from the results of the natural gamma ray log. sanyal et al. (1980) showed that lithology could be discriminated by analysing wire-line logs and making crossplots of especially the gamma-ray and neutron responses. some of the results are of special interest in the context of the present work on the lopra-1 logs. a positive correlation between flow thickness and mean velocity simply reflects increased proportions of massive basalt in thicker flows (planke 1994); velocity and density gradients at the top and bottom of a flow unit are largely independent of flow thickness (planke & flovenz 1996); goldberg et al. (1994) and kern & richter (1979) found that travel time correlates directly with porosity and showed that it also decreases with depth in sediments. planke (1994) found flows between 0.6 and 18.5 m thick while most sediment was less than 1 m thick. radioactive gamma ray and potassium logs are asymmetrical, having high values at flow tops, decrease gradually towards a central wide minimum and increase rapidly again near the flow base. principles of interpretation of the lopra-1 well wire-line data it is assumed that the principles and observations about the wire-line interpretation of subaerial basalt successions discussed in the previous section can be applied to the interpretation of the wire-line logging suite carried out in the lopra-1 well. the physical properties of a basalt flow are related primarily to active mechanical processes (i.e. growth through inflation, cooling history, the environment into which the flow was extruded, stress, faulting and whether it is a pahoehoe or aa flow type) and to secondary processes such as leaching and mineral growth. thus it is expected that, within subaerially extruded flows, the responses of the sonic, porosity, density, resistivity and calliper logs, all of which record physical properties, will follow a generalised cyclic pattern. on the other hand the geochemical properties reflect a combination of magma composition, primary mineralogy, weathering, leaching of various elements and growth of secondary minerals so the response from the gamma (sgr and cgr), thorium, potassium and uranium wire-line logs, all of which represent geochemical properties, will not necessarily show similar cyclic patterns. as a first approach, an idealised cross-section of a flow is envisaged as follows: a porous crust at the top of the flow grades into a massive core. a basal zone below the massive core appears abruptly and is located on top of the preceding tuff/sedimentary layer or, in case such a layer is absent, on top of the crust of the preceding flow. the porous crust represents a transition zone between the tuff/sediment layer above it and the massive core beneath it. its physical properties thus change gradually, represented by the readings of the physical logs changing gradually downwards from those observed at the very top of the flow or in the overlying tuff/sediment layer to the values of the underlying massive core. it is thus expected that values for the transit time of both the pand swaves, porosity and the diameter of the well will decrease downwards, whereas the density and resistivity will increase downwards. the interpretation of the geochemical logs is ambiguous because the crust could either be leached or enriched in elements compared to the original composition of the flow. the massive core is often found to be characterised by uniform values of the sonic, porosity, density, resistivity and calliper logs. more specifically, the core is recognised by the low transit time (high velocity) of the p-, sand stonely-waves, low values of porosity, the well diameter and drilling rate, but high values of density and resistivity. the chemical logs are expected to show a concentration geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1912 13 for the core and crust that more or less reflect the original chemical composition of the flow. the thin basal zone is recognised physically as a sharp transition above the underlying lava flow/tuff horizon. an abrupt change downwards of the log patterns is expected. values of transit time of p-, sand stonely-waves are expected to increase downwards (i.e. a downward fall in velocities) as are those of porosity and well diameter, but resistivity and density should decrease downwards. sediment deposited on top of a lava flow should appear on the petrophysical logs as follows: high values of transit times of p-, sand stonely-wave sonic logs (low velocity), high values of porosity, low values of density, resistivity, distinctly high values of the gamma ray, especially the spectral gamma ray responses (particularly for the compensated gamma ray and potassium curves) and enlarged diameter of the well. lopra-1 results the wire-line-logs are shown on enclosure 1. individual curves are arranged such that their visual impression can be correlated easily with the inferred lithology. inspection shows that while the basal zone of many flow units can be recognised, it is too thin to be treated statistically as has been done for the core, crust, compound flows, dolerite dykes and sediments. analysis of the basal zone has therefore been included in that of the core. calliper the calliper log shows that the diameter of the borehole varies cyclically in the range of 9–18″ in the upper part of the hole that was drilled in 1981 with an 8.5″ bit (enclosure 1). the diameter is within the interval 6.5″ to 13″ in the deepened hole drilled with a 6.5″ bit. the diameter tends to decrease with depth in the upper hole but it is close to being gauge only in a few places. the measured calliper was expected to have been close to gauge in size; that the final diameter of the hole is irregular is probably due to the method of drilling (nielsen et al. 1984). in the lower section drilled in 1996 with a different technique, the calliper is much closer to the bit size within cores (e.g. flow units 83, 86). the hole is narrowest in massive basalt and is usually wider in flow crusts and basal zones and is usually widest of all in sediment. in a few places, the largest diameter is found in porous crust rather than in overlying sediment, possibly indicating that the sediment is more compact or less friable than the lava crust. where the sediment beds are closely spaced, they do not seem to affect the condition of the hole, suggesting that the presence of nearby basalt stabilises the hole and that the sediment beds are rather competent (enclosure 1). in places, the massive cores of the flows show minor deviations from stable hole condition that might indicate that the core is fractured or not as massive as presumed (e.g. flow units 42, 46 and 58). p-, sand stonely-wave transit time the log curves in enclosure 1 have been plotted at the same scale but using different intersects, so that s-transit times are within the range 50–170 µsec/ft and p-transit times within 10–130 µsec/ft. plotting the curves in this way ensures that they plot on top of one another in massive cores where the hole is at or close to gauge as inferred from the combined interpretation of p-, sand stonely transit time, rhob and nphi in combination. the steady relationship between the two transit times probably indicates that the physical properties of the cores are fairly stable with a pto s-wave velocity ratio of 1.31. both pand s-wave transit time logs show a cyclic pattern (enclosure 1). the sediment beds are characterised by the highest transit time (70–100 µsec/ft sonic-p) followed by porous crusts and massive cores. maximum separation of the curves is a strong indicator of sediments, and intermediate separation is considered to represent crust. the velocity gradient in the upper part of flow units 24, 26, 28, 40, 41 and 42 seems to be independent of the thickness of the flow. transit times within the core units seem to be independent of the thickness of the flow (e.g. flow units 21, 29, 52, 59 and 79) (fig. 1). the stonely-wave transit time is a part of the full-wave sonic survey. attenuation of the stonely-waves occurs where fractures are open and may also be caused by changes in permeability. such attenuation will be more marked in hard formations where the acoustic contrast is greatest between the formation and mud-filled voids (rider 1996). the cyclic pattern outlined by the pand s-wave transit time logs is also shown on the stonely-wave curve, and the smallest amount of oscillation of the curve occurs in the dolerite dykes and in the core of the flow units (enclosure 1). density-porosity the bulk density and nphi porosity curves (measured in limestone porosity units) also show a cyclic nature. the geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1913 14 highest ‘porosity’ and lowest density values are recorded within the sediments, intermediate values in lava crusts and lowest ‘porosity’ and highest density values in the lava cores (enclosure 1). the curves have been plotted at comparable scales, which have been chosen so that they overlap within a typical massive basalt flow, which has been ascribed a density of 2.85–3.00 g/cm3 and a porosity (nphi) of 0.1–0.15. plotted in this way, a positive separation (where the porosity curve is located to the left of the density curve) is diagnostic of sediment or lava crust whereas no separation of the curves is diagnostic of lava core. negative separations occur within lava cores that have a lower than average porosity. resistivity logs the shallow (lls) and deep (lld) resistivity curves were obtained during the first drilling operations in the lopra-1 well. the log pattern shows a cyclic behaviour comparable to the other logs representing physical properties, reflecting the separation of flow units into crust, core and basal zones. the resistivity values are highest in the core of the flow units where the lld-log reads higher than the lls-log, showing an increase in resistivity away from the hole and into the formation. that could be due to either the formation water being fresher than the mud water and/or that the basalt becomes more massive away from the well. the values in the flow unit crusts are lower than those in the cores, and the separation between the curves is small. the composite flows have a behaviour similar to that of the crust. sediments have lowest values but the lld-values are still larger than those of the lls-log. dolerite dykes have the largest values of resistivity, and a large separation of the logs is found. geochemical logs five geochemical logs are presented: sgr and cgr gamma ray, potassium, thorium and uranium (enclosure 1). basalt that contains both consistently high and consistently low amounts of radioactive material is seen. the shifts from high to low radioactive basalt are remarkably abrupt, and there seems to be no gradual changes, especially for thorium, indicating that thorium represents the original magma (e.g. flow units 7–8, 12–13, 22–23, 23–24, 30– 31, 31–32, 46–47, 55–56, enclosure 1). in general terms, high thorium values are associated with high potassium and thus high cgr gamma ray, as well as high uranium values. there is a close relationship between the cgr gamma ray and potassium spikes and a less pronounced one between cgr gamma ray and uranium spikes. sediment is characterised by high localised potassium spikes, often recognised in the cgr gamma ray, that correlate to anomalies in the physical properties logs. in some flow unit flow unit 21 flow unit 29 flow unit 52 flow unit 59 flow unit 75 p-sonic cali_merged 140 40 8 23 nphi 0 0.5 2.2 rhob 3.2 cgr 0 30 thor 30 10 15 20 25 30 35 5 0 5 10 m fr om c ru st /c or e bo un da ry c ru st c or e c al ip er pso ni c fig. 1. variation with depth in five selected flow units. the bimodal behaviour within the flow units is found in the logs representing physical properties but not in the geochemical logs. variations with depth in the core are most pronounced for the nphi porosity and rhob density as compared to the psonic transit time and calliper. the psonic transit time appears rather constant. the cgr gamma-ray and thorium curves show the large variations present between different flow units. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1914 15 crusts, a decrease in potassium values with depth was found. lithofacies analysis the cyclicity of the wire-line log curves has been used to distinguish a number of lithological units including (1) sedimentary layers, (2) flow units of simple type composed of crust, core and basal zone, (3) groups of thin flow units presumably defining composite flows in the sense of walker (1991), (4) dolerite dykes. fifty-two units have been interpreted to be sedimentary beds, identified from a combined interpretation of the potassium, sgr and cgr gamma ray, sonic p-, sand stonely-waves, nphi porosity, rhob density, calliper and resistivity logs (enclosure 1). the sediment layers are recognised by gamma ray and potassium spikes relative to the overlying core (and basal zone) and the underlying crust. the gamma ray spike has to be larger than 25 gapi and be localised within a short thickness interval less than 3.0 m (98% of the sediment layers; fig. 2) and the pronounced potassium spike must contain more than 1% k. the uranium values are often high too, but are less diagnostic. in addition, the calliper shows a localised maximum deviation from the bit size, the sonic and nphi logs record high values whereas the density is low. the thicknesses of the sedimentary layers are generally small, and most of the bed thicknesses fall within the intervals 1.0– 1.5 m (38%) and 0.5–1.0 m (25%) (fig. 2). the total thickness of sediments is estimated to be 75.8 m as compared to 40 ± 4 m reported by nielsen et al. (1984). within the depth intervals displayed in enclosure 1, the nphiporosity and density curves often show a large negative separation. the nphi-porosity index is large, but the nphi-porosity often shows a higher value within the crustal part of the underlying flow unit. the density value is low, and generally the lowest value is found in the sediments rather than in the uppermost part of the underlying flow unit (e.g. flow units 12, 20, 29 and 61, enclosure 1). there are, however, examples where the density values are lower within flow units than within sediments. thus porosity and density alone are not diagnostic of sediment beds. the p-, sand stonely-wave transit times are high within small depth intervals. the shifted transit time curves of the pand s-waves show large negative separation, but the values are not diagnostic enough to separate sediment layers from the porous crust of the lava flow. the resistivity values are extremely low. hald & waagstein (1984) and waagstein et al. (1982) recognised a total of 36 sedimentary beds from analysis of cuttings sampled every 2 m and, among these, 32 are located below the casing of the well and thus within the 1996 log runs. it has been possible to establish a correlation between the sediment layers interpreted from cuttings and the 1996 log curves (enclosure 1), and 27 out of the 32 sedimentary beds can be confirmed by the criteria put forward here. thus 22 additional sedimentary layers have been identified within the shallower part of the lopra-1 well. a total of 87 flow units made up of crust, core and basal zone have been identified using a combination of p-, sand stonely-wave transit times, porosity and density values, resistivity values and the calliper log. from the top of each unit downwards to the core, the lava crust is indicated by a general gentle decrease in the sonic transit times combined with a decrease in porosity and increase in density and resistivity (enclosure 1). below this crustal zone, the log patterns are replaced by fairly constant low values of transit times and porosity combined with a high value of the density and resistivity, defining the lava core. the core is characterised by the calliper being closely in gauge and low values of the shifted sonic transit time curves, low and small oscillations of the stonely-transit time curve, no separation of the density and porosity curves, and high values of the resistivity values (enclosure 1). this fairly uniform behaviour of the log curves is abruptly different near the base of the unit where a pronounced increase in sonic transit time, porosity and calliper values combined with a pronounced decrease in density and resistivity values marks the basal zone. the thickness of the basal zone is generally small, and it has not been possible to differentiate it from the overlying core and the underlying sedimentary layer/crust in a consistent manner, so here the 0 5 10 15 20 fr eq ue nc y (% ) 1.0 2.0 3.00 4.0 sediment thickness in metres fig. 2. frequency diagram illustrating the distribution of sediment thicknesses interpreted in the present study. the sediment layers have been grouped into 0.5 m thickness intervals. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1915 16 basal zones have been included in the cores. the behaviour of the log curves representing the basal zone seems independent of whether or not a sediment layer is present beneath it. the individual flow units are generally reflected in the geochemical logs, primarily in the thorium curve and secondarily in the potassium curve, whereas the sgr, cgr gamma and uranium curves show unsystematic variations (e.g. flow units 6, 29, 39, 44, 51, 52, enclosure 1). the unsystematic variations do not allow an overall profile to be established as can be done with the logs representing the physical properties. it is therefore concluded that the geochemical behaviour of the flow units is largely related to the original composition of the magma and that different flow fields may be present. the thicknesses of the flow units have been compared to those interpreted by hald & waagstein (1984) and are found typically to be within the range 5–25 m (of which interval flow units of a thickness of 10–15 m represent 23%; fig. 3). a closer inspection of the curves through the crust and core parts of individual flow units reveals large fluctuations indicating that the physical properties are not constant in either the crust or the core. the nature of the variations in one crust is illustrated in a few examples; in flow unit 18 the porosity increases with depth and the density is rather constant, in flow unit 46 the porosity increases with depth and density decreases, in flow unit 20 the porosity decreases with depth and density increases. although the porosity and density figures for the crust differ, the relationship between the crust and core seems to be constant apart from those flow units where the crust and core are of similar thickness. the idealised core as a massive part of the flow unit characterised by constant low values of transit time, low values of porosity and high density and resistivity values is only true for a few flow units, e.g. 20 and 30, whereas different situations are found in e.g. flow units 13, 19, 26, 30, 36. most of the flow units penetrated by the lopra-1/1a well are characterised by the presence of porosity that may be vesicular zones or fractures within the core, though their regularity seems to indicate that they can be ascribed to vesicular zones. these zones are clearly reflected by the density and porosity curves and deflect the ssonic transit time curves to a larger degree than they do the p-sonic transit time curves. the vesicular zones within the core are not found at a consistent height above the base or below the top, but most of them seem to display some kind of symmetry. the variation from core to core is exemplified by e.g. flow units 26 and 40 which are most massive at the top and porous in the middle; the core of flow unit 41 is most massive in the upper part, and the vesicular zone is displaced towards the base as compared to the core of flow unit 40, the core of flow unit 58 is almost symmetrical with a porous zone in the middle, the cores of flow units 31 and 52 become more massive towards the base, and the core of flow unit 59 shows numerous vesicular zones. a minority of cores are massive in the middle and then are less massive both upwards and downwards (e.g. flow units 45 and 48). statistical analysis shows that when vesicular zones are present, the core is at least 70% thicker than the crust, whereas the core and crust are of equal thickness when no vesicular zones are indicated by the logs. five flow units (21, 29, 52, 59 and 75), representative of flow units of different thickness between 10 and 50 m, have been chosen to illustrate the relationship between flow thickness and mainly velocity. the thicknesses were correlated with the p-wave transit time, the calliper, the nphi-porosity, density, gamma ray (cgr) and the thorium content logs (fig. 1). it can be seen that different values of physical properties are recorded in the flow unit depending on whether crust or core is encountered. the flow units vary considerably in their geochemical expression, and the bimodal behaviour of the flow units is not seen. the crust is represented by high values of the calliper, transit time, nphi-porosity and low values of density. transit time in the core seems to be fairly constant and therefore independent of the thickness of the core and the flow unit. the variations in values recorded by the calliper seem independent of the thickness of the flow unit (fig. 1 and enclosure 1). the relationship between the flow unit thickness and nphi-porosity and density values of the core seems to suggest that thin flows exhibit high porosity and low density whereas thick flows show 0 5 10 15 20 fr eq ue nc y (% ) this study hald & waagstein (1984) 10 20 30 40 50 600 flow unit thickness in metres fig. 3. frequency diagram illustrating the distribution of flow unit thicknesses interpreted in the present study and in hald & waagstein (1984). the flow units have been grouped into 5 m intervals. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1916 17 crossplo flow-unit density/soni p/dept crossplo compoun flow density/soni p/dept crossplo crust density/soni p/deptcrossplo core density/son c /dep crossplo dolerite density/soni p/dept crossplo sediment density/soni cross-plot flow unit: density/p-sonic/depth cross-plot compound flow: density/p-sonic/depth cross-plot dolerite: density/p-sonic/depth cross-plot sediment: density/p-sonic/depth cross-plot core: density/p-sonic/depth cross-plot crust: density/p-sonic/depth 1.5 1.7 1.9 2.1 2.3 2.5 2.7 2.9 3.1 3.3 1.5 1.7 1.9 2.1 2.3 2.5 2.7 2.9 3.1 3.3 d en si ty 1.5 1.7 1.9 2.1 2.3 2.5 2.7 2.9 3.1 3.3 d en si ty d en si ty 1.5 1.7 1.9 2.1 2.3 2.5 2.7 2.9 3.1 3.3 d en si ty 1.5 1.7 1.9 2.1 2.3 2.5 2.7 2.9 3.1 3.3 d en si ty 1.5 1.7 1.9 2.1 2.3 2.5 2.7 2.9 3.1 3.3 d en si ty 50 70 90 110 130 50 70 90 110 130 50 70 90 110 130 p-sonic 50 70 90 110 130 p-sonic 50 70 90 110 130 p-sonic 50 70 90 110 130 p-sonic p-sonic p-sonic 2600 2200 1800 1400 1000 600 200 depth md fig. 4. cross-plot of inferred lithology: the cross-plots present the statistical behaviour of the flow units, cores, crusts, compound flows, dolerite dykes and sediments plotted for p-sonic transit time versus rhob density and depth. it is found that the clustering of data is enhanced by plotting the inferred lithology instead of the entire basalt column. 100 80 60 40 20 0 sonic (µs/ft) density flow unit massive porous compound dolerite sediment 40 140 40 140 40 140 40 140 40 140 40 140 1.50 3.30 1.50 3.301.55 3.151.50 3.301.50 3.301.55 3.15 avg: 55.127 min: 40.817 med: 52.79 max: 103.723 avg: 52.283 min: 40.817 med: 51.676 max: 92.807 avg: 65.0 min: 45.989 med: 64.4 max: 103.723 avg: 61.919 min: 46.422 med: 60.676 max: 125.3853 avg: 50.098 min: 40.227 med: 48.421 max: 90.512 avg: 71.606 min: 48.286 med: 70.878 max: 106.435 avg: 2.845 min: 1.554 med: 2.906 max: 3.151 avg: 2.900 min: 1.566 med: 2.931 max: 3.151 avg: 2.658 min: 1.554 med: 2.664 max: 3.151 avg: 2.739 min: 2.033 med: 2.757 max: 3.011 avg: 2.900 min: 1.510 med: 2.946 max: 3.035 avg: 2.582 min: 1.529 med: 2.582 max: 2.949 100 80 60 40 20 0 fr eq ue nc y (% ) fr eq ue nc y (% ) fig. 5. histogram and statistical values of the six cross-plots presenting the inferred lithology for p-sonic transit time versus density and depth shown in fig. 4. the average, median, minimum and maximum values are shown. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1917 18 low porosity and high density (fig. 1) but enclosure 1 shows that this relation is too simple. the term ‘compound flow’ is used for intervals where the cyclicity of the logs representing physical properties suggests strongly that flow units are present, but the individual parts of the flows are poorly defined and log readings do not reach the values normally ascribed to crust or core (enclosure 1). eighteen compound flow intervals have been recognised. dolerite dykes have previously been encountered at two levels in the lopra-1 well (hald & waagstein 1984). the dykes are represented by constant values of gamma ray, thorium and potassium lower than found in flow units, rather high density and the lowest values of porosity and transit time combined with the highest values of the resistivity measured in the well (enclosure 1). the calliper shows that pronounced caving occurred in the upper dyke, but no such caving occurred in the lower one. the dyke intervals were expected to show rather constant values of the log curves representing the physical properties but, as is seen on enclosure 1, a cyclic appearance of those logs is seen. this cyclic appearance is most obviously seen in the lower dolerite dyke where five symmetric intervals are revealed by the rhob, nphi, lls and lld logs. this suggests that the lower dolerite dyke consists of five thinner dykes. statistics the lithological division of the volcanic sequence into flow units (consisting of porous crust and massive core), compound flows, dolerite dykes and sediments makes possible a statistical analysis of the properties of the various lithologies found in the lopra-1 well. the data obtained from especially the physical properties logs as exemplified by the values of p-, sand stonely-wave, porosity, density and resistivity logs (table 3) show a much reduced scatter after being separated into the lithological divisions. a strong correlation, most pronounced for the core and to a lesser extent for the crust and sediment beds, is also found between pand s-wave transit times, pand s-wave transit times versus density, pand s-wave transit times versus porosity, and density versus porosity. examples are presented in figs 4 and 5 where six cross-plots with associated histograms and statistical parameters show the benefit 200 800 1400 2000 40 60 90 120 140 200 800 1400 2000 40 60 90 120 140 200 800 1400 2000 40 60 90 120 140 200 800 1400 2000 2600 2600 d ep th ( m ) 40 60 90 120 140 flow units composite flows dolerites sediments cores crusts 200 800 1400 2000 40 60 90 120 140 800 1400 2000 40 60 90 120 140 2600 200 2600 2600 2600 d ep th ( m ) d ep th ( m ) d ep th ( m ) d ep th ( m ) d ep th ( m ) sonic (µs/ft) sonic (µs/ft) sonic (µs/ft) sonic (µs/ft) sonic (µs/ft) sonic (µs/ft) fig. 6. cross-plot of p-sonic transit time versus depth for the inferred lithology. the figure shows that the core and sediment lithologies especially show a pronounced trend of decreasing transit time values with depth whereas crust and compound flows are ambiguous. the amount of data representing dolerite dykes is insufficient to describe their relationship with depth. the flow units show a strong increase in transit time with depth but this is due to the core part of the flow unit. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1918 19 of dividing the geological column into the suggested lithological units. the relationship between the physical properties of the lithological units and depth has been examined and a correlation is observed. the degree of correlation is strongest for the flow units (due to the core values), core, sediment/ tuff layers, dolerite dykes and to a lesser extent for crust and compound flows. p-sonic transit time curves are illustrated versus depth for the different lithologies as an example (fig. 6). there it is seen that especially the sonic transit time increases with depth for the flow units and a first order regression line has been estimated as p-sonic = 58.1370–0.0022 × depth measured in metres or p-sonic = 58.0934–0.0007 × depth when depth is measured in feet. likewise a first order regression line estimated for the cores is p-sonic = 54.0165–0.0012 × depth in metres and p-sonic = 54.7769–0.0011 × depth for depths measured in feet. this observation is in accordance with general findings made by kern & richter (1979). the sediment/tuff layers also show a significant correlation with depth and in general terms the crust does too, whereas the compound flows show no systematic variation. a depth relation for dolerite dykes is not justified because of their limited numbers. the potassium and cgr gamma ray logs show a strong response to the sediment beds, but there is otherwise no correlation between the geochemical logs and the individual flow units. this probably indicates that the geochemical values are related to the original composition of the flow units, which varies strongly, suggesting that one magma type may represent more than one flow unit. this can be illustrated by analysing the thorium content in relation to flow units and compound flows as defined by the petrophysical logs. the intervals covering the dolerite dykes are not included in this part of the study as they are reported to be later intrusions (hald & waagstein 1984). the thorium log data have been block averaged into three groups according to their thorium content (group 1 consisting of 72 flow and compound flow units with a thorium content of 0.0–1.0 ppm; group 2 containing 21 flow and compound flow units with 1.0–2.0 ppm th; group 3 with 11 flow and compound flow units representing a content of 2.0–3.0 ppm th). this results in 36 groups, tentatively interpreted as flow fields (enclosure 1). some flow fields are constructed of a number of flows and compound flow units, but 57% of the flow fields constitute one flow or compound flow unit (fig. 7). the individual flow fields are separated by sediment/tuff layers in 70% of cases. since this study is an analysis of the basalt column based primarily on a lithological division from petrophysical logs, further statistical treatment of the geochemical logs is outside its scope and probably should be done by dividing the lava flows into those with high, medium and low radioactivity. discussion the flood basalt pile in the interval 200–2486 m depth in the lopra-1 well has been divided into lithological units based on the wire-line logging (enclosure 1). eighty-six flow units consisting of porous crust and massive core (including basal zones), 18 compound flows and two dolerite dykes have been recognised using petrophysical logs (transit time of p-, sand stonely-waves, nphi-porosity, density, resistivity and calliper). fifty-two sedimentary/tuff layers have been identified mainly by potassium, cgs and sgr gamma logs supplemented by the petrophysical logs. thirty-seven flow fields have been suggested based on the thorium log in combination with the petrophysical logs. the flow units have been subdivided into three parts on criteria proposed by various workers based on field evidence or wire-line logs from basalt covered regions (e.g. planke 1994; planke & flovenz 1996; waagstein 1999). the lopra-1 results reveal that overall the crust, core and basal zones are similar to other basalt covered areas reported using wire-line log studies, but there are some differences. to describe the crusts as showing a downward increase in density and resistivity, a decrease in porosity and pand s-wave transit times is too simple. the downward increase in density values is commonly replaced by zones of decreasing values, probably reflecting the presence of vesicular or fractured zones. likewise the previously reported behaviour of the core as being rather uniform although most massive in its middle part (e.g. planke 1994) is far from true for the flows of the lopra-1 well. number of flow units and compound flows in flow field 11 5 1 2 3 5 64 11 14 20 fr eq ue nc y (% ) flow field with 0–1 ppm th flow field with 1–2 ppm th flow field with 2–3 ppm th fig. 7. frequency diagram illustrating the distribution of the number of flow units or compound flows versus thorium content included in the inferred flow fields. the flow units and compound flows have been grouped into three depending on their thorium content. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1919 20 two types of core are commonly present, one type with a ‘massive’ central part where the logs indicate that only the central part is truly massive. this type, which represents 12% of the flow units, is generally thin, and the crust is of the same thickness as the core. the other type is characterised by varying numbers of zones of high porosity/low density whose relative depth within the core also varies. in these flow units the thickness of the core is at least 70% of the total thickness of the flow unit, and the flow units themselves are of different thicknesses. this zonation is interpreted as vesicular or perhaps occasionally fractured zones. the presence of vesicular zones has been reported by mcmillan et al. (1987) and walker (1989). walker (1989) described the vesicular zonation from spongy pahoehoe lava in hawaii and mentioned that aa-type lava can also contain vesicular zones. this raises the question whether cores of pahoehoe lavas can be distinguished from those of aa-type lava. most papers distinguish aa from pahoehoe using surface morphology and perhaps the vertical thickness of the basal zone (waagstein 1999). unfortunately surface morphology is not revealed by the present logging suite, and the vertical thicknesses of the basal zone are too small to be treated on their own in the present study. however, hald & waagstein (1984) reported that the majority of basalt from the shallower part of the lopra-1 well was aa-type lava and that pahoehoe lava in the field seems to have a similar crust and core relationship. it is therefore tentatively suggested that flow units with vesicles may represent aa-type flows, and that the flow units where the crust and core are of the same thickness and the core is massive are pahoehoe type. however only a few studies have dealt with the question of how a flow unit is formed in relation to its distribution of vesicles and cooling history based on cross-sections. the present study has identified 86 flow units and 18 compound flows compared with the 87 reported by waagstein et al. (1982) and hald & waagstein (1984). flow thicknesses of 10–15 m (23%) are most common followed by thicknesses of 15–20 m (18%) (fig. 3). thus the flow unit thicknesses reported here are slightly higher than those of hald & waagstein (1984). five flow units of different thickness – tentatively interpreted as aa-type – were chosen by using the transit time of the p-wave, calliper, nphi-porosity, density, thorium and gamma ray (cgr) logs to analyse the thickness of the flow units (fig. 1). the petrophysical logs show clearly that different values are ascribed to the crust and core. the expected overall trend of the flow unit in relation to the petrophysical logs is valid. the p-wave transit time in the core is rather constant and not dependent on the thickness of the core and thus of the flow unit. in contrast to the transit time, the nphi-porosity and density varies, and in the flow units chosen the density seems to increase and the porosity decrease with thickness of the core and thus of the flow unit. that the transit time is independent of the thickness of the core and thus of the flow unit is somewhat different from the results presented by planke (1994) from a seaward dipping reflector sequence. the behaviour of the radioactive gamma ray and potassium logs in lopra-1 also differs as they do not show the cyclicity found in the seaward dipping reflectors studied by planke (1994). whether this indicates different depositional settings, lava types, magma type, weathering conditions or the fact that fresh water is encountered in the lopra-1 well is uncertain. eighteen compound flows, characterised by similar values of the logs reflecting the physical properties as the crust, have been interpreted. however a cyclic behaviour similar to that in the flow units can be recognised within the compound flows although the values of the pand swave transit time, density, porosity and resistivity are lower than in the flow units. it is suggested that the compound flows represent a number of thin flows or the toes of flow units. two dolerite dykes were encountered and are characterised by the lowest values of sgr and cgr gamma ray, pand s-wave transit time and porosity combined with the highest level of resistivity, high values of density and a constant level of the stonely-wave transit time curve. the sonic pand s-transit time logs and the density, porosity and resistivity curves show a cyclic behaviour resembling the pattern of a flow unit, and it is speculated that the dyke intervals represent a group of thinner dykes. fifty-two sedimentary layers have been detected by the combination of potassium, cgr and sgr gamma log, transit time logs of p-, sand stonely-waves, nphi-porosity, density, resistivity and calliper logs. the sediments have been identified on the basis of a locally high spike of the potassium log value in combination with a gamma ray spike larger than 25 gapi, high values of transit time of p-, sand stonely-waves, high values of the nphiporosity and low values of the density and resistivity. the number of sedimentary layers is greater than reported by hald & waagstein (1984) and nielsen et al. (1984) who reported 27 sediment layers based on the combination of cuttings and geophysical logs acquired in 1981. of the 27 sediment layers identified by hald & waagstein (1984), the recently acquired geophysical logs used in the present study have confirmed 22 of them. the sedimentary layers are thought to have formed during periods of quiescent volcanism. in that case, at least 51 periods of quiescence occurred in the interval from geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1920 21 184 to 2484 m depth in the lopra area. mostly a single flow unit was extruded between each pause, but up to eight flows are found within one interval. the enlarged number of sediment layers indicates that the time intervals between eruptions taking place were shorter than previously thought and thus the extrusion rate of basalt was faster than previously assumed. the amount of basalt produced between the periods of quiescence, measured by the thickness of the flow unit, differs markedly at different depths in the lopra-1 well. this may indicate that the thickness of basalt produced by a single eruption differed, but to confirm this the location and extension of each eruption needs to be known. the geochemical logs, especially the thorium log, show numerous fluctuations indicating the presence of some highly radioactive basalt flows, similar to those reported in the deccan trap area by buckley & oliver (1990). to try to systematise these fluctuations three intervals have been defined (0–1 ppm, 1–2 ppm and 2–3 ppm), and the column was subdivided at the same time into smaller units that appear to correlate with the flow units and sedimentary beds, and on this basis it is suggested that a minimum of 36 units exists. this suggests that 36 different compositions of magma have been present during the time represented by the lopra-1 well, and tentatively it is suggested that a minimum of 36 flow fields exists. table 2 lists results of other studies of subaerially extruded basalts whereas table 3 contains the statistical values of the lopra-1 well wire-line log data from this study. the lopra-1 results, which are more detailed due to the subdivision into the suggested lithology units, the large thickness of basalt and the extensive logging suite, are in overall agreement with the range of results obtained from the literature survey. some scattering in measured values exists; however, the amount of data is not sufficient to conclude whether it is the primary volcanic setting or postdepositional processes that are responsible for these differences. at a depth of approximately 2417 m, the cyclic nature of the wire-line log curves representing the physical properties of the basalt terminates abruptly, and an increase in transit times and porosity combined with a decrease in density signal a transition into the non-subaerially extruded basalt sequence below this level. the transition generates a negative seismic reflection from this level. it is concluded that the negative reflection interpreted from the vsp by kiørboe & pedersen (1995) corresponds to the base of the subaerially extruded basalt. the total known thickness of the flood basalt on the faroe islands can therefore be increased by 250 m. conclusions the study shows that it is possible to use wire-line logging combined with results from work carried out in flood basalt regions in other parts of the world to divide the subaerial plateau basalts drilled in the lopra-1 well into a number of lithological units. fifty-two sedimentary layers have been identified using geochemical logs supplemented by the logs representing physical properties. eighty-six basaltic flow units (consisting of a vesicular crust, a massive or partly vesicular core and a thin basal zone), 18 compound flows and two dolerite dykes have been identified using the logs representing physical properties supplemented by the geochemical logs. tentative aa and pahoehoe lava types are suggested, among which the aa lava is the most common. the geochemical logs appear to respond to the composition of the magma rather than reflecting individual flow units, and it is suggested that the lopra-1 well penetrated a minimum of 36 flow fields. the present subdivision of the lithology is strongly supported by statistical analysis of the logged physical parameters and shows that a pronounced reduction in data scatter is obtained when the basalts are studied as individual units. thus the study presents consistent and detailed information about the lower series of the faroe plateau basalts that has led both to new results and results in greater detail than was previous known, which is of benefit in geophysical and geological studies of basalt and pre-basalt geology. acknowledgements thanks are due to the lopra consortium and the geological survey of denmark and greenland for access to the wire-line log data. thanks are due to dr. h. delius for assistance in correcting the resistivity logs. thanks are due the reviewers professor dr. d. goldberg and professor dr. p.k. harvey for valuable and constructive suggestions that have improved the content. j.a. chalmers is thanked for valuable suggestions that improved the language. landmark is acknowledged for the university grant issued to the geological institute, university of copenhagen. references berlitz, r. et al. 1988: log interpretation in igneous and metamorphic rocks with five case studies. technical review 36, 30–47. berthelsen, o., noe-nygaard, a. & rasmussen, j. (eds) 1984: the deep drilling project 1980–1981 in the faeroe islands. annales geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1921 22 societatis scientiarum faeroensis, supplementum ix, 158 pp. tórshavn: føroya fróðskaparfelag. buckley, d.k. & oliver, d. 1990: geophysical logging of water exploration boreholes in the deccan traps, central india. in: hurst, a., lovell, m.a & morton, a.c. (eds): geological application of wire-line logs. geological society special publication (london) 48, 153–161. cashman, k.v. & kauahikaua, j.p. 1997: reevaluation of vesicle distribution in basaltic lava flows. geology 25, 419–422. christensen, n. 1996: poisson’s ratio and crustal seismology. journal of geophysical research 101(b2), 3139–3156. goldberg, d.s., reynolds, d.j., williams, c.f., witte, w.k., olsen, p.e. & kent, d.v. 1994: well logging results from the newark rift basin coring project. scientific drilling 4, 267–279. hald, n. & waagstein, r. 1984: lithology and chemistry of a 2 km-sequence of lower tertiary tholeiitic lavas drilled on suðuroy, faeroe islands (lopra-1). in: berthelsen, o., noe-nygaard, a. & rasmussen, j. (eds): the deep drilling project 1980– 1981 in the faeroe islands. annales societatis scientiarum faeroensis, supplementum ix, 15–38. tórshavn: føroya fróðskaparfelag. kern, h. & richter, a. 1979: compressional and shear wave velocities at high temperature and confining pressure in basalts from the faeroe islands. tectonophysics 54, 231–252. kiørboe, l.v. & pedersen, s.a. 1995: vsp seismic experiment in the faeroes. in: scrutton, r.a. et al. (eds): the tectonics, sedimentation and palaeoceanography of the north atlantic region. geological society special publication (london) 90, 111–123. lund, j. 1983: biostratigraphy of interbasaltic coals from the faeroe islands. in: bott, m.h.p. et al. (eds): structure and development of the greenland–scotland ridge. nato conference series iv: marine science, 417–423. new york/london: plenum press. mcmillan, k., randal, c.w. & long p.e. 1987: two-stage vesiculation in the cohassett flow of the grande ronde basalt, southcentral washington. geology 15, 809–812. nielsen, p.h., stefánsson, v. & tulinius, h. 1984: geophysical logs from lopra-1 and vestmanna-1. in: berthelsen, o., noenygaard, a. & rasmussen, j. (eds): the deep drilling project 1980–1981 in the faeroe islands. annales societatis scientiarum faeroensis, supplementum ix, 115–135. tórshavn: føroya fróðskaparfelag. planke, s. 1994: geophysical response of flood basalts from analysis of wire-line logs. ocean drilling program site 642, vøring volcanic margin. journal of geophysical research 99(b5), 9279– 9296. planke, s. & flovenz, o.g. 1996: seismic properties of flood basalt (extended abstract), 6 pp. geophysics for lithology predictions. kristiansand: norwegian petroleum society. rasmussen, j. & noe-nygaard, a. 1969: beskrivelse til geologisk kort over færøerne i målestok 1:50 000. danmarks geologiske undersøgelse 1. række 24, 370 pp. + map vol. 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manuscript received 20 june 2000; revision accepted 11 september 2001. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1922 geological survey of denmark and greenland bulletin 9 1 geological survey of denmark and greenland bulletin 9 · 2006 scientific results from the deepened lopra-1 borehole, faroe islands edited by james a. chalmers and regin waagstein geological survey of denmark and greenland danish ministry of the environment 2 geological survey of denmark and greenland bulletin 9 keywords faroe islands, palaeogene basalts, lopra-1/1a borehole. cover maersk rig 81 on the location of the lopra-1/1a wells at suðuroy, faroe islands, in august 1996 shortly after the start of drilling. photo: regin waagstein. frontispiece: facing page maersk rig 81 on the location of the lopra-1/1a wells at suðuroy, faroe islands, in august 1996 shortly after the start of drilling. photo: regin waagstein. chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, geological institute, university of aarhus; minik rosing, geological museum, university of copenhagen; finn surlyk, geological institute, university of copenhagen scientific editors of this volume: james a. chalmers and regin waagstein editorial secretaries: esben w. glendal and birgit eriksen referees: t. bidstrup (denmark), d. bird (usa), r. burwood (uk), b. christenson (new zealand), a. förster (germany), d.s goldberg (usa), p. harvey (uk), s. jakobsson (iceland), r. løvlie (norway), h. micheelsen (norway), j.d.a. piper (uk), m.a. rooney (argentina), d. tarling (uk), g. van graas (norway), r. white (uk), m. worthington (uk) and an anonymous referee (denmark) illustrators: eva melskens and kristian rasmussen digital photographic work: benny m. schark graphic production: knud gr@phic consult, odense, denmark printers: schultz grafisk, albertslund, denmark receipt/acceptance dates of manuscripts: see end of individual articles printed: 31 may 2006 issn 1604-8156 isbn 87-7871-179-7 geological survey of denmark and greenland bulletin the series geological survey of denmark and greenland bulletin replaces geology of denmark survey bulletin and geology of greenland survey bulletin. citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 9, 156 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk or geografforlaget aps rugårdsvej 55, dk-5000 odense c, denmark phone: +45 63 44 16 83, fax: +45 63 44 16 97, e-mail: go@geografforlaget.dk © danmarks og grønlands geologiske undersøgelse (geus), 2006 3 4 contents introduction m.v. heinesen, a. rosenkrands larsen and k. sørensen . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5 wire-line log-based stratigraphy of flood basalts from the lopra-1/1a well, faroe islands l.o. boldreel . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7 borehole seismic studies of a volcanic succession from the lopra-1/1a borehole in the faroe islands, northern north atlantic p. christie, i. gollifer and d. cowper . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23 magnetic logs from the lopra-1/1a and vestmanna-1 wells, faroe islands n. abrahamsen and r. waagstein . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 41 palaeomagnetic results from the lopra-1/1a re-entry well, faroe islands n. abrahamsen . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 51 petroleum geochemistry of the deepened lopra-1/1a re-entry well, faroe islands j.a. bojesen-koefoed and h.p. nytoft . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 67 hydrocarbon gases in palaeogene volcanic rocks from the lopra-1/1a well, faroe islands t. laier . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 79 thermal structure of the deep lopra-1/1a borehole in the faroe islands n. balling, n. breiner and r. waagstein . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 91 mineralogical and thermodynamic constraints on palaeogene palaeotemperature conditions during low-grade metamorphism of basaltic lavas recovered from the lopra-1/1a deep hole, faroe islands w.e. glassley . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 109 a reconnaissance study of fluid inclusions in fracture-filling quartz and calcite from the lopra-1/1a well, faroe islands j. konnerup-madsen . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 119 the regional distribution of zeolites in the basalts of the faroe islands and the significance of zeolites as palaeotemperature indicators o. jørgensen . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 123 composite log from the lopra-1/1a well, faroe islands r. waagstein . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . in pocket inside back cover 5 introduction martin v. heinesen, arne rosenkrands larsen and kai sørensen systematic preparation of the legal framework to govern full-scale hydrocarbon exploration in the faroe islands was initiated in 1992, after the danish government and the faroese home rule government had agreed on the future administrative regime with respect to minerals, including hydrocarbons, in the subsurface of the faroese area. at about the same time, significant hydrocarbon discoveries were made in the uk offshore area west of the shetland islands and close to the uk–faroese border that led to production starting from the foinaven and schiehallion fields later in the decade. for these reasons the petroleum exploration community became interested in the faroese area. the stratigraphic section exposed in the islands consists predominantly of a 3000 m thick series of mainly tholeiitic basaltic lavas of late paleocene age, divided informally into a lower, a middle and an upper series. until the early 1990s, the only deep wells that had been drilled in the faroese area were the vestmanna-1 well, drilled in 1980 in the town of vestmanna on streymoy, and the lopra-1 well drilled on the southernmost part of suðuroy in 1981, both scientific wells funded mainly by the faroese government with a contributing grant from the carlsberg foundation of copenhagen (organising committee 1984). vestmanna-1 was a 660 m deep, fully cored, slim well that penetrated the lowermost part of the middle series and the uppermost part of the lower series. lopra-1 was located near where the stratigraphically lowest basalts are exposed. it was originally drilled to a depth of 2178 m and penetrated approximately 2 km of formerly unknown strata, all of which proved to be a continuation of the subaerially exposed lower basalt series. the well was left open, plugged only by a 10″ gate valve. the uppermost part of the well was cased with a 16 m conductor pipe with a diameter of 14″ and a 95/8″ casing down to 190 m, whereas the 8½″ diameter main part of the well remained uncased down to terminal depth (td). although the lopra-1 well did not reach the base of the volcanic pile beneath the faroe islands, it revealed some information of significance to the hydrocarbon exploration potential of the area. no strong evidence was seen of any sedimentary section nor hydrocarbon generation in the well, but analysis of tiny amounts of methane-rich gases and oil isolated from the water in the well indicated an origin from marine organic matter beneath the basalt (jacobsen & laier 1984). subsequent seismic experiments, including different vertical seismic profiling (vsp) experiments, indicated a lithological change only c. two hundred metres beneath the td of the well (kiørboe & petersen 1995). after consultation with a number of oil companies, in 1995 the faroese petroleum administration made an arrangement with the danish state oil company (dong) to act as operator and coordinator of a joint industry project to deepen the lopra-1 well. the 19 oil companies that participated in the project are listed in table 1.the petroleum administration was assisted by the faroese geological survey and the geological survey of denmark and greenland (geus). dansk operatørselskab i-s (danop) acted as consultant to dong with respect to the drilling operation and the mærsk rig 81 drilling rig was used. geus and the faroese geological survey were responsible for the well site geology and the subsequent geological evaluation of the drilling results. the primary objective of the deepened well was to ob© geus, 2006. geological survey of denmark and greenland bulletin 9, 5–6. available at: www.geus.dk/publications/bull amerada hess ltd. amoco dk exploration co. arco british ltd. bhp petroleum inc. bp exploration operation co. ltd. british gas exploration & production ltd. chevron europe ltd. dansk olieog gasproduktion a/s (dong) deminex uk oil & gas ltd. enterprise oil exploration ltd. esso exploration. & production uk ltd. fina research s.a. lasmo (ulx) ltd. mobil north sea ltd. norsk hydro a.s. phillips petroleum co. uk ltd. saga petroleum international a.s. shell exploration b.v. statoil efterforskning og produktion a.s. table 1. the nineteen oil companies that participated in the project 6 tain lithological and stratigraphic information about the deepest parts of the faroe plateau basalt series and their substratum and to acquire information in relation to the hydrocarbon prospectivity of the area. secondary objectives included: (a) determination of the age and depth to the base of the basalt series, (b) characterisation of the pre-basaltic lithologies, (c) characterisation of the physical properties of the basalt series and their substratum to improve interpretation of regional seismic, gravity and magnetic data and (d) thermal maturity determination of interbasaltic and pre-basaltic organic matter, as well as determination of possible sources for oil and gas shows in the well. deepening of the well started on 13 july 1996. it reached its total depth of 3565 m on 3 november 1996. after final wireline logging, it was plugged and abandoned and the rig was released on 13 november 1996. the well penetrated 213 m of subaerially extruded basalt flows of the lower basalt series, a series of pillow lavas 45 m thick and a series of pillow lava debris 41 m thick. below that was a thick series of volcanic tuffs with some intra-volcanic sandstone and claystone stringers. the only conventional core cut in the well was a 1.5 m (recovered) core from the lowermost basalt flow. drilling through the base of the volcanics was not achieved. this book presents a number of research studies carried out in the years following completion of the well. although no direct information was obtained about prevolcanic rocks, the deepening of the lopra-1 well has documented the presence on and under the faroe islands of the thickest known composite section through the north atlantic palaeogene basalt plateau. references jacobsen, o.s. & laier, t. 1984: analysis of gas and water samples from the vestmanna-1 and lopra-1 wells, faeroe islands. in: berthelsen, o., noe-nyegaard, a. & rasmussen, j. (eds): the deep drilling project 1980–1981 in the faeroe islands. annales societatis scientarium færoensis, supplementum ix, 149–155. tórshavn: føroya fróðskaparfelag. kiørboe, l. & petersen, s.a. 1995: seismic investigation of the faeroe basalts and their substratum. geological society special publication (london) 90, 111–112. organizing committee 1984: preface. in: berthelsen, o., noenyegaard, a. & rasmussen, j. (eds): the deep drilling project 1980–1981 in the faeroe islands. annales societatis scientarium færoensis, supplementum ix, 5 only. tórshavn: føroya fróðskaparfelag. _____________________________________________________________________________________________________________________________ m.v.h., jardfrødisavnid (jfs) – faroese geological survey, p.o. box 3169, fo-110 tórshavn, faroe islands. present address: jardfeingi – faroese earth and energy directorate, p.o.box 3059, fo-110 tórshavn, faroe islands. e-mail: martin.v.heinesen@jardfeingi.fo a.r.l., dansk olie og naturgas (dong) , agern alle 24–26, dk-2970 hørsholm, denmark. k.s., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. geological survey of denmark and greenland bulletin 17, 2009, 45-48 the european union’s water frame work directive aims to achieve a ‘good’ ecological status for groundwater bodies, for groundwater-dependent terrestrial ecosystems, and for aquatic surface water bodies by the year 2015. in denmark, this goal will most likely not be fulfilled within such a short time frame due to the current poor ecological condition of danish lakes (søndergaard et al. 2008). however, public concern about the protection of aquatic environments has increased, and so has interest in improving lake water quality by reducing nutrient loading. effective and sustainable lake restoration and conservation depend on the ability to (1) point out sensitive catchment areas for the lake, (2) estimate its total water and nutrient budgets and (3) relate observed differences in seepage rates to the abundance and distribution of macrophytes in the lake and to the topography and land-use of the surrounding terrain. in seepage lakes, i.e. lakes without inlets or outlets, the influence of the surrounding terrain, regional hydrogeology and lake geometry on the overall lake water budget has been studied in some detail (krabbenhoft et al. 1990; anderson & cheng 1993; cheng & anderson, 1994; kratz et al. 1997; water budget of skærsø, a lake in south-east jylland, denmark: exchange between groundwater and lake water bertel nilsson, peter engesgaard, jacob kidmose, sachin karan, majken caroline looms and mette cristine schou frandsen © geus, 2009. geological survey of denmark and greenland bulletin 17, 45–48. available at: www.geus.dk/publications/bull 45 fig. 1. a: location of skærsø east of the main stationary line in jylland. b: lake area in summer (dark blue) and winter (light blue). c: map of the lake skærsø area showing the groundwater level of the regional aquifer and the water level of skærsø. black dots show the locations of wells used to construct the map. 9°16.5’e skærsø 55°35’n contours (1 m interval) lake in summer lake in winter 500 m 50 m 4849 44 45 4647 66–67 m cb a 150 m 9°16.5’e 43 m 55°35’n weichselian deposits saalian deposits outwash plains marine deposits aeolian deposits freshwater deposits 50 km main stationary line rosa_2008:rosa-2008 01/07/09 15:48 side 45 winter 1999; townley & trefry 2000). however, little effort has been made to understand and quantify how riparian zones (wetlands) surrounding lakes may control the water flow and nutrient transport to the lakes. although groundwater inflow to seepage lakes is suspected to be smaller than inflow from drainage ditches, it may still account for a significant nutrient influx. this paper focuses on field work carried out 2007 and 2008 at skærsø, a lake in the upper part of the kolding å catchment area in south-east jylland (fig. 1a). field studies at skærsø have shown that seepage can vary on different scales in both space and time. the purpose of this study is to link measurements of lake seepage rates and lake precipitation or evaporation to the catchment hydrogeology. the project is conducted by the centre for lake restoration, which includes participants from university of southern denmark, national en vironmental research institute, university of copenhagen and geological survey of denmark and greenland. towards a lake typology the centre for lake restoration is developing a typological classification of danish lakes using a multidisciplinary ap proach that integrates interactions between groundwater and lake water. the classification is based on geological, hydrological, hydrogeological, geomorphological, botanical and chemical aspects. the botanical part focuses on plant indicator species in the lakes, and the chemical part addresses water and sediment chemistry. two or three main lake types are currently distinguished on the basis of geological, hydrological, hydrogeological and geomorphological criteria. new lake types will probably be defined in the coming years as biolo gical and chemical indicators are also included in the classification. location and setting of skærsø skærsø is located a few kilometres east of the main stationary line that formed during the last glacial maximum around 20 000 years ago. the lake is situated in the upper part of the 46 medium k low k high k 4 77 44 17 19 39 medium k low k high k k = hydraulic conductivity a b unsaturated zone lake lake fig. 2. a: preliminary model of the water balance of skærsø. grey arrows show influx, black arrows loss. numbers in per cent. b: conceptual hydrogeological model of a perched lake disconnected from the underlying aquifer. a a sand sand gyttjagyttja sandsand multiplesmultiples sandsand sand gyttja sand multiples 50 m 8 0 2 4 6 d ep th ( m ) sand bbb fig. 3. a: mapping of lake sediments using reflection ground penetrating radar (gpr). the radar equipment is contained in the grey rubber boat. b: an example of a gpr reflection profile with a preliminary interpretation. rosa_2008:rosa-2008 01/07/09 15:48 side 46 kolding å catchment area (fig. 1a) that is dominated by glacial and glaciofluvial deposits. around skærsø a 2–7 m thick sand layer overlies an approximately 10 m thick clayey till that is underlain by a regional sandy aquifer. the riparian zone around skærsø has a width of 30–50 m and a thickness of 0.1–0.5 m, and consists of organic-rich fine sand and silt. the water table of the lake varies seasonally about 0.5 m, and in the winter or early spring the riparian zone around the lake is flooded (fig. 1b). during flooding, organic particles and solutes are transported to the lake and cause a significant decrease in the transparency of the water. below skærsø, the groundwater table of the regional sand aquifer is 48–49 m above sea level, whereas the surface of the lake is 66–67 m above sea level. this difference indicates that there is a hydraulic connection between the lake and the regional aquifer (see below and fig. 1c). the area of skærsø is approximately 16 ha , and its average water depth is 1.5 m, with a maximum depth of about 8 m. it is characterised as a mesotrophic, low-alkaline, clear-water lake. until the 1980s, skærsø was a clear-water heath lake, dominated by submerged macrophytes including lobelia dortmanna that grew to a depth of 1.5 m. at the end of the 1980s, skærsø suffered from organic-rich, acidic, unclear water, resulting in almost complete disappearance of the submerged macrophytes. after almost 19 years of unsuccessful attempts at lake restoration, the conditions in skærsø are still deteriorating. a recent study showed that changes in light conditions in the lake are not as strongly coupled to changes in the external nutrient loading as expected (frandsen & stæhr 2007). changes in the hydrology of the riparian zone have led to a seasonal flush of dissolved coloured organic matter into the lake which causes decreased water transparency, decreased light penetration and release of nutrients. however, the mechanisms that affect water transparency require further investigation. hydrogeological model the hydrogeological model of the lake is characterised by the setting of shore and wetland (the riparian zone) with an upper, local groundwater aquifer in sediments of low to moderate permeability showing low hydraulic gradients. the deeper part of the lake is probably connected to a lower permeable layer with high hydraulic gradients (fig. 1c). because the surface of skærsø is almost 20 m above the hydraulic head of the regional sand aquifer, we expect leakage of water from the lake through its bottom into the underlying regional aquifer. the water supply to skærsø is dominated by precipitation (77%) and inflow from the catchment area (19%). a limited portion (4%) comes from the shallow, local groundwater aquifer. water flux from the lake is dominated by outflow via outlets (44%; mainly an artificial ditch) and evaporation (39%). we suggest that the remaining 17% is lost by leakage to the underlying aquifer (fig. 2a). thus skærsø can be conceptualised as a perched lake, a lake which is disconnected from the underlying aquifer by an unsaturated zone between the lake bottom and the aquifer (fig. 2b). ground penetrating radar in order to investigate the structure and thickness of the sedi ments below the lake, a ground penetrating radar survey was carried out (fig. 3a). thick layers of organic-rich sediments such as lake mud (gyttja) or peat generally reduce ground water seepage, whereas sandy sediments can promote inter action between groundwater and lake water (fig. 3b). ground penetrating radar can help to identify areas of potential groundwater seepage, which may then be verified using more traditional point measurements such as coring. the appli cation of ground penetrating radar to map lake sediments is new in denmark, and the interpretation of the radar profiles is still uncertain and needs to be checked against data from coring. however, according to our preliminary interpretation of the radar data, sandy sediments are widespread below the lake, whereas gyttja is probably restricted to a small area (fig. 3b). multidisciplinary approach to estimate the groundwater flux seepage to or from skærsø to the upper, local groundwater aquifer was measured using various tracers (heat, stable isotopes), direct measurements of water flux, as well as nutrients sampled from piezometer transects. the combined use of different tracers and methods at different scales provides a good understanding of the physical, chemical and biological behaviour of the entire lake. figure 4 shows three types of field equipment that were used to quantify the flux from the local aquifer to skærsø, and the estimated groundwater fluxes (specific discharges) to skærsø are summarised in table 1. the flux values estimated by the three different methods are not consistent. the highest estimated fluxes (using the darcy 47 method flux specific discharge (m/sec) seepage meter q seep 10–7 – 3 × 10–7 temperature q t 10–8 – 10–7 darcy1 q d 10–7 – 10–6 table 1. groundwater discharge into skærsø 1 q d = k × i, where k is the saturated hydraulic conductivity of the lake-shore sediment (i.e. fine sand and silt in the upper 2 m) equiv alent to k = 10–4 – 10–5 m/sec, and i is the hydraulic gradient measured by a potentiomanometer to about ± 0.01. rosa_2008:rosa-2008 01/07/09 15:48 side 47 (transect) method) are 10–100 times higher than the lowest estimated fluxes (using the temperature method). despite this difference, the data indicate a low inflow from the shallow groundwater aquifer and, as mentioned above, we suggest that 4% of the inflow to the lake comes from the shallow aquifer, based on the average value obtained by the applied methods. we have no data on leakage through the lake bottom to the underlying aquifer, but we consider this process likely and suggest a flux value of 17% based on the difference between the other fluxes. however, this figure is highly uncertain. especially the figure for the evaporation from the lake may be underestimated if investigations of other lakes in jylland are considered. nevertheless, we conclude that exchange between groundwater and lake water plays an important role in the water budget of lake skærsø. acknowledgement villum kann rasmussen fonden is gratefully acknowledged for funding the centre for lake restoration (clear), which is a villum kann rasmussen centre of excellence. references anderson, m.p. & cheng, x. 1993: longand short-term transience in a groundwater/lake system in wisconsin, usa. journal of hydrology 145, 1–18. cheng, x. & anderson, m.p. 1994: simulating the influence of lake position on groundwater fluxes. water resources research 30, 2041–2049. frandsen, m.c.s. & stæhr, a.p. 2007: differentieret analyse af lysmiljøet i skærsø, 24 pp. unpublished report, freshwater biological laboratory, university of copenhagen, hillerød, denmark. krabbenhoft, d.p., bowser, c.j., anderson, m.p. & valley, j.w. 1990: estimating groundwater exchange with lakes. 2. calibration of a threedimensional, solute transport model to stable isotope plume. water resources research 26, 2455–2462. kratz, t.k., webster, k.e., bowser, c.j., magnuson, j.j. & benson, b.j. 1997: the influence of landscape position on lakes in northern wisconsin. freshwater biology 37, 209–217. lee, d.r. 1977: a device for measuring seepage flux in lakes and estuaries. limnology and oceanography 22, 140–147. nilsson, b., engesgaard, p., kidmose, j. & karan, s. 2008: groundwater – lake exchange at lake skærsø in western denmark. proceedings of 36th iah congress, october 2008, toyama, japan. integrating groundwater science and human well-being. extended abstract on cd-rom. schmidt, c., bayer-raich, m. & schirmer, m. 2005: characterization of spatial heterogeneity of groundwater – stream water interactions using multiple depth streambed temperature measurements at the reach scale. hydrology and earth system sciences 10, 849–859. søndergaard, m., liboriussen, l., pedersen, a.e. & jeppesen, e. 2008: lake restoration by fish removal: shortand long-term effects in 36 danish lakes. ecosystems 11, 1291–1305. townley, l.r. & trefry, m.g. 2000: surface water – groundwater interaction near shallow circular lakes: flow geometry in three dimensions. water resources research 36, 935–949. winter, t.c. 1999: relation of streams, lakes, and wetlands to groundwater flow systems. hydrogeology journal 7, 28–45. winter, t.c., labaugh, j.w. & rosenberry, d.o. 1988: the design and use of a potentiomanometer for direct measurement of differences in hydraulic head between groundwater and surface water. limnology and oceanography 33, 1209–1214. authors’ addresses b.n., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: bn@geus.dk p.e., j.k., s.k. & m.c.l., department of geography and geology, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. m.c.s.f., freshwater biological laboratory, biological institute, university of copenhagen, helsingørsgade 51, dk-3400 hillerød, denmark. 48 a b c fig. 4. equipment used for field determination of groundwater seepage. a: modified 1 m long potentiomanometer used to measure the water pressure in the subsurface (winter et al. 1988). b: multi-level, 1.25 m long temperature probe used to measure the temperature at multiple levels in the sediment (schmidt et al. 2005). c: a steel drum seepage meter, 0.5 m in dia meter, used to measure the flow of ground water through the lake bottom (lee 1977). from nilsson et al. (2008). rosa_2008:rosa-2008 01/07/09 15:48 side 48 geological survey of denmark and greenland bulletin 20, 2010, 103–106 103 in the following we describe the result of the titan project, an interactive web application (titan) developed at the geological survey of denmark and greenland (geus) together with dupont titanium technologies. the main aim of titan is to make computer-controlled scanning electron microscopy (ccsem) data, generated at geus, available via the internet. in brief, ccsem is a method automatically to detect particles with a scanning electron microscope (sem), and based on computer-controlled imagery to measure the chemistry and grain morphology of each particle in a given sample (knudsen et al. 2005; bernstein et al. 2008); keulen et al. 2008. titan makes data available on-line so that the user can interact with the data sets and analyse them using a web browser. in addition to ccsem data, titan contains a global database of titanium deposits and various reports. the web application is customised, such that the functionality and amount of data available for a given user depend on the privileges of that user. data a prerequisite for making titan available on the internet is that the data are stored in a relational database allowing for fast querying and retrieval. a detailed description of the data model is not given here, but we will make a short overview of the different data types and how the data are inserted and managed in the database. the data are stored in an oracle 10g database that is housed and maintained by geus (tulstrup 2004). ccsem data ccsem data are uploaded to the jupiter database via a desktop program developed with the delphi software tool at geus (fig. 1). typical ccsem samples consist of roughly one thousand grains. for each grain c. 30 analytical parameters are measured including elemental composition, grain size and grain shape. all parameters are stored in the database including the energy dispersive x-ray spectrum for each grain (from which the elemental composition of the grain is determined) as well as backscatter-mode sem images of the sample. the number of measurements currently stored in the database exceeds 54 million (fig. 2), and the number continues to grow at a rate of c. 15 million per year. once data are uploaded, each grain is classified according to a mineral classification scheme that is implemented using oracle database pl/sql procedures. the mineral classification scheme is based on the measured concentration of elements, and different schemes are available depending on the type of sample. for example, there is a mineral classification scheme for heavy mineral concentrates, while another scheme is applied for soil samples. titanium deposits and reports data on titanium deposits can be entered directly into the database through titan, and at present it contains detailed information about more than 600 deposits around the world. interactive web analysis and presentation of computer controlled scanning electron microscopy data peter riisager, nynke keulen, uffe larsen, roger k. mclimans, christian knudsen and jørgen tulstrup © geus, 2010. geological survey of denmark and greenland bulletin 20, 103–106. open access: www.geus.dk/publications/bull fig. 1. entry of sample metadata and uploading of ccsem data to the database using a desktop programme. the energy dispersive x-ray spectrum for each grain and backscatter-mode sem images of the sample are also uploaded. 104104 reports and report metadata are uploaded to the database via the web application. the option to create, read, update, and delete deposits and reports is restricted to certain users and managed by strict user control. web interface the web application is based on open-source software, mainly java, and is running in an open-source jboss application server housed and maintained by geus. access to titan requires login with username and password. once logged in, the system assigns the user a specific role with specific priviliges allowing certain analytical functions. the user control is also extended to the individual ccsem samples so that each group of users can see and interact with only a restricted subset of ccsem samples. the ccsem data can be accessed using an interactive geological map web page (fig. 3) or by alphanumeric search criteria. when a sample is found, various recalculated data for that sample are available in table formats, and back-scatter mode sem images or other files, such as outcrop pictures related to the sample, can be viewed or downloaded. if users want to work with the data off-line, data can be exported to microsoft excel format or a customised pdf report can be generated. most importantly, the user has the option to visualise and interact with the data using various plots. for example, the elemental composition of all the mineral grains in a given sample can be presented in a scatter plot (fig. 4), where the user defines the elements assigned to the x and y axes and which minerals should be plotted. the user can select which minerals to be plotted for analysis and comparison of apparent grain sizes (fig. 5). this gives the user quick and easy access to make the desired analyses and plots. the on-line access has the added benefit that users around the world always have live access to the database, meaning that as soon as new data are uploaded the users can start interacting with them. 2002 2003 2004 2005 2006 2007 2008 2009 2010 10000000 20000000 30000000 40000000 50000000 year 0 n u m b er o f an al ys es fig. 2. number of measurements in the titan database as a function of time. currently the database contains c. 54 million measurements. fig. 3. the titan web application includes an interactive map that allows the user to search for ccsem samples based on their geographic sampling coordinates. clicking on a sample on the map takes the user to other screen images where one can interact with the actual data (figs 4, 5). 105 examples of applications titan has successfully delivered data to scientists working at geus as well as various external end-users. in the following we will give some short examples of the use of titan. tracing kimberlite-indicator minerals for diamond prospection the elemental compositions of megacrystal and xenocrystal phases within kimberlitic rocks are used as an important diamond exploration tool. preliminary studies have demonstrated the excellent potential of ccsem to determine the elemental composition of minor elements in these minerals in a reliable and more cost-efficient manner than conventional electron microprobe analysis (keulen et al. 2009). exploration and ore deposit evaluation mineralogical characterisation of sediments is a prerequisite for exploration and exploitation of valuable sediment occurrences, such as heavy mineral deposits. with titan we have characterised individual mineral particles in heavy mineral sands, with the aim of detecting heavy mineral ore deposits (knudsen et al. 2005; bernstein et al. 2008). soil sample mineralogy for the cement industry cement manufacture causes emission of large quantities of airborne pollutants including greenhouse gases. with titan we have studied the raw materials used in cement production to optimise the performance of the grinding mill and the sintering process, with the ultimate aim to lower energy consumption at high temperatures (keulen et al. 2008). sediment provenance studies for oil exploration and sedimentary basin analysis sediment provenance studies are an important aspect in the evaluation of possible sandstone reservoirs. titan has served as an indispensable tool to determine source, compositional fig. 4. analysis of ccsem data with an interactive scatter plot using titan. the user defines the elements assigned to the x and y axes, and which minerals are plotted. 106106 variation and sedimentary pathways of several sedimentary deposits relevant for oil exploration (knudsen et al. 2005; bernstein et al. 2008). concluding remarks earth science is becoming an increasingly more quantitative science with new insights, more often than not, derived from detailed measurements. the titan project represents an example where advanced geoscientific instrumentation has allowed us to design projects that were not previously feasible, thereby opening new scientific research areas (e.g. bernstein et al. 2008; keulen et al. 2008, 2009). it is, however, important to underline that this progress does not only include the scientific instruments and analyses themselves, it also represents a major challenge to store and analyse the huge quantities of data generated (fig. 2). in this paper we have demonstrated that information technology plays a central role in facilitating the storage, retrieval and analysis of large geoscientific datasets. storing and distributing geological data are important aspects of national geological surveys, and it is important that geus follows in the footsteps of other national surveys in order to take full advantage of the new opportunities offered by the revolution of information technology. the titan project demonstrates that geus has the in-house expertise to develop technologies that address the full life cycle of geoscientific data from the raw instrument output, to store data securely in a database, analyse and reduce data, and finally distribute the data providing on-line and interactive access to the data. the know-how from the titan project is currently being applied to other similar projects at geus. finally, we would like to point out that besides the scientific and societal rationale of titan, it continues to be used by commercial partners. in fact, the application itself has almost exclusively been financed by industrial partners. references bernstein, s., frei, d., mclimans, r.k., knudsen, c. & vasudev, v.n. 2008: application of ccsem to heavy mineral deposits: source of high-ti ilmenite sand deposits of south kerala beaches, sw india. journal of geochemical exploration 96, 25–42. keulen, n., frei, d., bernstein, s., hutchison, m.t., knudsen, c. & jensen, l. 2008: fully automated analysis of grain chemistry, size and morphology by ccsem: examples from cement production and diamond exploration. geological survey of denmark and greenland bulletin 15, 93–96. keulen, n., hutchison, m.t. & frei, d. 2009: computer-controlled scanning electron microscopy: a fast and reliable tool for diamond prospecting. journal of geochemical exploration 103, 1–5. knudsen, c., frei, d., rasmussen, t., rasmussen, e.s. & mclimans, r. 2005: new methods in provenance studies based on heavy minerals: an example from miocene sands in jylland, denmark. geological survey of denmark and greenland bulletin 7, 29–32. tulstrup, j. 2004: environmental data and the internet: openness and digital data management. geological survey of denmark and greenland bulletin 4, 45–48. authors’ addresses p.r., n.k., u.l., c.k. & j.t., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: pri@geus.dk r.k.mcl., dupont titanium technologies, experimental station, e352/217, route 141 and henry clay, wilmington, de 19808, usa. fig. 5. analysis and comparison of the apparent grain size distributions of the various mineral fractions in the sample using titan. the plot is interactive, allowing the user to choose which minerals are plotted. geological survey of denmark and greenland bulletin 4, 2003, pp 97-100 97 the vietnam petroleum institute (vpi) and the geological survey of denmark and greenland (geus) have carried out a programme of geoscientific research and institutional capacity building since 1995. it has included geoscientific projects focused on assessment of the hydrocarbon potential of selected sedimentary basins, technology transfer and inhouse training at vpi in hanoi and at geus in copenhagen. co-operation is continuing within the framework of a new, long-term project that aims to strengthen research capacity in vietnam within the fields of basin analysis and modelling. history of co-operation in 1993 ccop-representatives (co-ordinating committee for geoscience programmes in east and southeast asia) visited geus to investigate the possibility of co-operation within research capacity development in the ccop-countries. the danish energy authority (dea) subsequently approved a project proposal, and the first joint project between vpi and geus was initiated in 1995. the commencement of geoscientific studies showed an urgent need for interpretation tools at vpi, and modern geological and geophysical interpretation facilities were installed at vpi in 1996 with support from the danish international development assistance (danida). following the successful completion of the first geoscientific project, dea approved funding for a second phase. the projects were accompanied by a series of ccop workshops held at vpi, and ‘on-thejob-training’ (ojt) programmes at geus for vpi geoscientists (nielsen et al. 2003). the two first geoscientific projects were aimed at an analysis of the hydrocarbon potential of the song hong basin (fig. 1). seismic interpretation, field work, stratigraphic evaluation, and assessment of results from the few exploration wells drilled in the basin were carried out (dien et al. 1997, 1998, 1999; andersen et al. 1998, 1999; nielsen et al. 1999). seismic facies analyses together with investigations of thermally immature potential source rocks at localities in the song hong basin where inversion structures expose oil-prone oligocene lacustrine mudstones and thin coals, led to the identification of areas for future hydrocarbon exploration (petersen et al. 2001, in press). the implementation of these initial projects revealed a fundamental need for enhancement of the research skills of vpi geoscientists as part of an institutional development programme. an application for a long-term geoscientific research capacity building project was submitted to danida’s enreca project in the year 2000. basin analysis and assessment of hydrocarbon potential the enreca project was approved in august 2001. the main aims were to undertake an integrated analysis and modelling of sedimentary basins in vietnam, to assess their hydrocarbon potential and to establish a geoscientific background geological survey of denmark and greenland bulletin 4, 97–100 (2004) © geus, 2004 petroleum potential of sedimentary basins in vietnam: long-term geoscientific co-operation with the vietnam petroleum institute lars henrik nielsen and ioannis abatzis fig. 1. regional map showing the prospective vietnamese cenozoic offshore basins. on which decisions concerning exploration and exploitation of hydrocarbons could be based. the studies are carried out by geoscientists from vpi and geus, with support from scientists at the hanoi university of mining and geology, vietnam and the university of copenhagen, denmark. the first phase of the project, focusing on the cenozoic phu khanh basin offshore central vietnam (fig. 1), will last for three years; long-term co-operation can be extended in up to four phases. one of the principal short-term objectives is to provide a research-driven assessment of the hydrocarbon potential of the basin. activities initiated during the first phase of the project include seismic interpretation, field work, drilling of a 500 m deep core-well, geochemical and petrographic analyses, the initiation of five vietnamese and four danish m.sc. studies, and several training courses. seismic interpretation the phu khanh basin is almost unexplored; it is covered only by an open seismic grid and no wells have yet been drilled. the cenozoic basin-fill overlies strongly faulted older sedimentary rocks and granitic basement that crop out along the present-day coastline. the basement is cut by deep, listric, eastward-dipping faults formed during rifting, which probably began during the palaeogene. the basin-fill consists of a syn-rift succession separated from the overlying post-rift succession by an unconformity that may be related to the opening of the south china sea. on seismic profiles the syn-rift wedges are characterised by continuous reflectors interpreted to represent lacustrine sediments, i.e. potential source rocks, whereas transparent zones are interpreted to represent fluvial sands (fig. 2). the middle miocene to recent post-rift succession can be divided into eight seismic sequences forming an eastwardthickening succession. the lower three sequences are thin, and their limited distribution is in part determined by structural uplift, whereas the more widely distributed overlying five sequences show distinct basinward prograding shelf-edge clinoforms. analyses of the seismic facies and geometry of the sequences suggest a variety of sandy depositional environment, such as fluvial channels, shorefaces and basin floor fans that may all form potential reservoirs for hydrocarbons generated from syn-rift lacustrine source rocks. high amplitude reflectors over structural highs are interpreted as miocene carbonate build-ups that can also form potential reservoirs. oil seeps oil seeps occurring at the basin margin have long been known from onshore coastal outcrops, but their nature has remained controversial. however, geochemical analyses of petroleum impregnation in cracks and fractures in weathered cretaceous granites, tar mats on the beach, oil on vegetation, and mud with oil from shrimp-farm basins, have shown the presence of biodegraded petroleum, with biological markers pointing to a marly source rock with some terrestrial input (bojesen-koefoed et al. 2003, 2004). the oil contains 30norhopanes characteristic of carbonate-derived petroleum, as well as oleanane, a marker of contributions from angiosperm higher land plants. the presence of petroleum in narrow fissures in freshly cleaved granite testifies to the indigenous 98 fig. 2. seismic profile of the phu khan basin. the coastal oil seeps are located immediately west of the profile and indicate an updip migration path of hydrocarbons generated from a possible kitchen in the phu khanh basin. for location, see fig. 1. nature of the oil, and an origin from oil spills is therefore unlikely. samples from shrimp basins show the presence of similar oil, although more altered by biodegradation. one sample collected from a tar mat on the beach shows characteristics similar to those of cenozoic lacustrine oils found elsewhere in vietnam, for instance in the song hong basin. the close similarity of these oils suggests that pods of prolific lacustrine petroleum source rocks may be regionally distributed. rift-lake successions and source rocks in order to investigate the formation and hydrocarbon potential of rift-lake systems associated with transtensional faulting, outcrop studies were carried out in the song ba trough supplemented by studies of a continuous core from the 500 m deep enreca-1 well. the song ba trough was formed by transtensional faulting in the onshore continuation of the tuy hoa shear zone that marks the southern boundary of the phu khanh basin (fig. 1). the outcrops expose coarsegrained sandstones and conglomerates interbedded with lacustrine mudstones and scattered coal beds. the core-well encountered 480 m of sediments, beginning with thin fluvial sandstones overlying granite. the sandstones are overlain by lacustrine mudstones interbedded with thin sandstones, whereas fluvial sandstones dominate the upper part of the core. the lacustrine mudstones contain organic matter, commonly exceeding 5 wt%, and many samples have a hydrogen index (hi) of more than 300 mg hc/g toc (total organic carbon), occasionally exceeding 700 mg hc/g toc (fig. 3). the organic matter is mainly composed of fluorescent amorphous organic matter, alginite and liptodetrinite, corresponding to type i kerogen, and may be an excellent source for oil. the results suggest that oil-prone mudstones were primarily formed in periods when the sedimentation rate was outpaced by formation of new accommodation space. significance for future exploration activities and research interpretation of the seismic profiles suggests the presence of possible reservoirs and deep-seated kitchens in the phu khanh basin. one possible kitchen is situated down-dip along a likely migration route to the oil seeps (fig. 2). thus, there is evidence that the common petroleum system found in south-east asian basins, which rely on syn-rift lacustrine mudstones and coals, is also active in the phu khanh basin. the investigated rift-lake succession in the onshore song ba trough may thus be a valid analogue for older and deeply buried syn-rift sequences in the phu khanh basin. 99 fig. 3. core log of a selected interval of the enreca-1 core-well, which penetrated a 500 m thick rift-succession of lacustrine mudstones. a significant amount of preserved organic matter indicates an excellent source rock for oil generation. the core-well is located in the song ba trough, see fig. 1. hc, hydrocarbon; hi, hydrogen index; toc, total organic carbon. 100 acknowledgements the energy research programme of dea, the royal danish embassies in hanoi and bangkok, danida’s fellowship center and danida are thanked for their financial support. vpi and petrovietnam are acknowledged for their institutional support and permission to publish the results. references andersen, c., tiem, p.v., mathiesen, a. & nielsen, l.h. 1998: some new thermal maturity modelling results using the yükler 1 d software and seismic facies mapping in the northern part of the song hong basin. in: toan, t.n., quy, n.h. & ngoc, n.v. (eds): proceedings of con ference on vietnam petroleum institute 20 years development and prospects. hanoi, vietnam, may 1998, 273–284. andersen, c., mathiesen, a., nielsen, l.h., tiem, p.v. & dien, p.t. 1999: petroleum systems in the northern part of the song hong basin, gulf of tonkin – vietnam. proceedings from international conference on tectonics, stratigraphy and petroleum and mineral systems of palawan, borneo and surrounding areas, palawan island, philippines, 29 november – 3 december 1999 (cd-rom). bojesen-koefoed, j.a., nytoft, h.p., dau, n.t., ha, n.t.b., hien, l.v., quy, n.h., nielsen, l.h. & petersen, h.i. 2003: geochemical characteristics of seep oils from dam thi nai (qui nhon), central vietnam – implications for exploration in the offshore phu khanh basin. 21st international meeting on organic geochemistry, krakov, poland, 8–12 september, 2002. abstracts 2, 193–194. bojesen-koefoed, j.a., nytoft, h.p., dau, n.t., ha, n.t.b., hien, l.v., quy, n.h., nielsen, l.h. & petersen, h.i. 2004: seep oils from dam thi nai (qui nhon), central vietnam; the enreca-project reports promising implications for the future exploration in the offshore phu khanh basin. in: jarupongsakul, t. & saito, y. (eds): 5th international conference on asian marine geology. igcp475 deltamap/ apn megadelta, bangkok, thailand, 13–18 january. abstracts p. 25 only. dien, p.t., nielsen, l.h., andersen, c. & nhuan, d.v. 1997: late mesozoic to cenozoic basin development along the north-west margin of the east vietnam sea. petrovietnam review 4, 5–10. dien, p.t., nielsen, l.h., andersen, a., tiem, p.v. & nhuan, d.v. 1998: late mesozoic – cenozoic events along the north-west margin of the east vietnam sea. in: toan, t.n., quy, n.h. & ngoc, n.v. (eds): proceedings of conference on vietnam petroleum institute 20 years development and prospects. hanoi, vietnam, may 1998, 125–131. dien, p.t., quy, n.h., tiem, p.v., tai, p.s., andersen, c. & nielsen l.h. 1999: basin analysis and petroleum system of the song hong basin. in: hiep, n. et al. (eds): geology and petroleum in vietnam, 44–67. hanoi, vietnam: youth publishing. nielsen, l.h., mathiesen, a., bidstrup, t., vejbæk, o.v., dien p.t. & tiem, p.v. 1999: modelling of hydrocarbon generation in the cenozoic song hong basin, vietnam; a highly prospective basin. journal of asian earth sciences 17, 269–294. nielsen, l.h., abatzis, i., petersen, h.i., bojesen-koefoed, j. & nytoft, h.p. 2003: geo-scientific co-operation between vietnam petroleum institute and geological survey of denmark and greenland, 1995–2002. results, lessons learned, and views for the future. petrovietnam review 2, 32–48. petersen, h.i., andersen, c., anh, p.h., bojesen-koefoed, j.a., nielsen, l.h, nytoft, h.p., rosenberg, p. & thanh, l. 2001: petroleum potential of oligocene lacustrine mudstones and coals at dong ho, vietnam – an outcrop analogue to terrestrial source rocks in the greater song hong basin. journal of asian earth sciences 19, 135–154. petersen, h.i., nytoft, h.p. & nielsen, l.h. in press: characterisation of oil and potential source rocks in the northeastern song hong basin, vietnam: indications of a lacustrine-coal sourced petroleum system. organic geochemistry. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: lhn@geus.dk << /ascii85encodepages false /allowtransparency false /autopositionepsfiles true /autorotatepages /all /binding /left /calgrayprofile (dot gain 20%) /calrgbprofile (srgb iec61966-2.1) /calcmykprofile (u.s. web coated \050swop\051 v2) /srgbprofile (srgb iec61966-2.1) /cannotembedfontpolicy /warning /compatibilitylevel 1.4 /compressobjects /tags /compresspages true /convertimagestoindexed true /passthroughjpegimages true /createjdffile false /createjobticket false /defaultrenderingintent /default /detectblends true /colorconversionstrategy /leavecolorunchanged /dothumbnails false /embedallfonts true /embedjoboptions true /dscreportinglevel 0 /syntheticboldness 1.00 /emitdscwarnings 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<feff4f7f752890194e9b8a2d5b9a5efa7acb76840020005000440046002065874ef65305542b8f039ad876845f7150cf89e367905ea6ff0c4fbf65bc63d066075217537054c18cea3002005000440046002065874ef653ef4ee54f7f75280020004100630072006f0062006100740020548c002000520065006100640065007200200035002e0030002053ca66f465b07248672c4f86958b555f3002> >> >> setdistillerparams << /hwresolution [2400 2400] /pagesize [595.000 842.000] >> setpagedevice geological survey of denmark and greenland bulletin 17, 2009, 9-12 a moderately strong earthquake struck southern sweden 5 km south-west of the town of sjöbo, 60 km east of malmö, in the early morning at 6:20 a.m. local time on 16 december 2008. the epicentre was located in skåne, a region that is known for its extremely low seismicity, and its location was determined to be 55.5°n and 13.6°e with an uncertainty of about 6 km. a depth of 9 km with an uncertainty of 3 km was obtained from teleseismic observations at the yellowknife seismic array, usa. since waveform data from the swedish national seismic network are not yet available, depth estimation using local stations has so far not been attempted. dur ing the period 1970–2008, only three small earthquakes were detected in the region; the largest measured 2.8 on the local richter scale. to our knowledge none of these previous earthquakes were felt by people. the historical archives dating back to 1375 show that 14 other earthquakes have been felt in the area. the largest of these, recorded in 1894, was felt over an area of 7300 km2 and had an epicentre 50 km east of the 16 december 2008 earthquake (scandinavian earthquake archive 2003). the activity in southern sweden is similar to that of northern sjælland and north-western jylland, and confirms the low seismicity of the region (gregersen et al. 1991). even though the earthquake was unusual, the event was not totally unexpected. northern europe is un der constant pressure from the mid-atlantic ridge, and the resulting stress is released in small to moderate size earthquakes. the local stress field is further modified by postglacial rebound. earthquakes in denmark and southern sweden occur where there are weaknesses or faults in the subsurface. in some cases the earthquakes occur where there are no mapped faults (gregersen et al. 1996). shaking from the earthquake was felt widely in denmark. strong shaking and low-frequency earthquake sounds frightened many people and caused them to abruptly leave their houses. authorities such as the police and the geological survey of denmark and greenland (geus) were flooded with phone calls and e-mails from concerned citizens. fortunately the earthquake did not cause any real damage, and no one was in danger at any time. macroseismic effects in denmark the basic physical earthquake parameters are determined from instrumental recordings. however, the effects of the earthquake on people, buildings, and nature require human inter© geus, 2009. geological survey of denmark and greenland bulletin 17, 9–12. available at: www.geus.dk/publications/bull earthquake in southern sweden wakes up denmark on 16 december 2008 peter h. voss, tine b. larsen, lars ottemöller and søren gregersen 9 6˚e 8˚e 10˚e 12˚e 14˚e 16˚e 58˚n 57˚n 56˚n 200 km 55˚n fig. 1. map showing the epicentre of the 16 december 2008 earthquake (yellow circle) and the locations of the first 3000 macro seismic reports received (red circles) from denmark. known faults are also shown. rosa_2008:rosa-2008 01/07/09 15:47 side 9 action to be delineated. immediately following the earthquake, geus requested input from the public through the news media and our own web page. our earthquake questionnaire was placed on www.geus.dk and a paper copy was mailed to those who had no internet access or who felt uncomfortable about reporting electronically. after a few days, geus had received more than 3000 earthquake reports (fig. 1), and after a month the number exceeded 4000 reports. this is a large increase compared to the last widely felt earthquake in den mark in 2001, where we received little more than 400 reports (larsen et al. 2008) and the one in 1985 in kattegat, which was almost as large as the 16 december 2008 earthquake, where we received a total of around 500 reports from sweden and denmark. in denmark alone the area where the earthquake was felt covers about 50 000 km2 (fig. 1). earthquakes in scandinavia are felt over much larger areas than in california, where c.f. richter worked when he invented the richter magnitude scale. thus the attenuation in scandinavia is lower than that in california. a comparison of local observations of 25 earthquakes in different regions and their magnitude suggests that the affected area this time will exceed 100 000 km2 (fig. 2). reports from the population are an important tool to assess the impact of an earthquake and to identify locations that are particularly vulnerable to ground shaking. this is still of great interest as earthquakes can be compared to historic events predating the era of instrumental recording. this information is used in seismic hazard studies, for example in connection with large construction projects. the geus questionnaire comprises the address of the observer, information about the building in which the observer was located and detailed information about shaking and other effects caused by the earthquake. shaking is strongly amplified in tall buildings, and the euro pean macroseismic scale (grünthal et al. 1998) recommends that shaking observations above the 5th floor are not taken into account when determining the intensity of an earthquake. a detailed study of the 4000 current reports has not yet been completed, but the reports will provide valuable shake information especially for sjælland. in many locations the earthquake was felt strongly. some examples, with locations shown on fig. 3, are: copenhagen: “walls, closet and bed were shaking”, gilleleje: “the chair i was sitting in was shaking”, humlebæk: “it was as if the bed was shaken – it felt very terrifying”, ålborg: “it felt as if someone was shaking my bed gently”, kettinge: “as when a large truck is just outside”. some reports claim that the earthquake caused a road in helsingør to crack. inspection of the road confirms the presence of cracks, but it is not obvious that the cracks opened during the earthquake. new asphalt was added to the road a year earlier, and the small cracks are located in the weak zone where new asphalt overlaps with old. the road is exposed to heavy traffic and the cracks could have opened at any time. they do not look any different from other asphalt cracks. the euro pean-mediterranean seismological centre has also received information (a photo) of cracked asphalt in sweden, but it has not been confirmed that these cracks formed during the earthquake. 10 6 10 00 2 3 4 magnitude 5 10 00 0 00 10 0 00 0 10 0 00 a re a (k m 2 ) intensity 3 uk noshel nor kads fig. 2. perceptibility areas for 25 selected earthquakes (area of intensity 3) correlated with their magnitudes. the red squares are for the shield region in kaliningrad, denmark and sweden (kads) compared to the 16 december 2008 earthquake. these areas are larger than those in norway (nor), on the norwegian continental shelf (noshel) and in the united kingdom (uk). the 16 december earthquake (preliminary result shown by a circle) seems to confirm the trend. sjöbo helsingør malmö jylland sjælland copenhagen humlebæk kettinge ålborg holbæk skåne skagerrak kattegat øresund gilleleje fig. 3. map of denmark and south-western sweden showing the location of place names mentioned in the text. rosa_2008:rosa-2008 01/07/09 15:47 side 10 earthquake magnitude determining the magnitude of an earthquake is not straightforward. many different definitions of magnitude are in use at seismological data centres around the world. they measure fundamentally different parameters in the seismograms, all with the intention to reproduce c.f. richter’s original magnitude measure from 1935 (richter 1935), which is in widespread use for public information and in engineering in earthquake regions. the earthquake hazards program at the united states geological survey (usgs) calculates up to six different magnitude values for each registered earthquake. at geus, two different magnitudes are calculated: our own computation of local magnitude (ml), which should give the classical richter number (richter 1935) for local earthquakes, and the moment magnitude (hanks & kanamori 1979), which reflects the size of the fault area and the slip during the earthquake. at geus, the ml is calculated from the maximum amplitude of the earthquake signal on the vertical component of the seismograph (usually the (surface) lg wave) taking into account the distance between the earthquake and the seismograph station. when the lg wave is observed at several seismographs, we calculate the ml as the average of the separate ml values. it requires little data to calculate and it can be calculated shortly after the earthquake. when the data quality is low, as is often the case with weak, local earthquakes, the ml is the only magnitude that can be calculated. the calculation of ml builds on experience of how effectively the shaking propagates in the area. the variations in the ml value calculated at different seismograph stations can be large. following the earthquake on 16 december 2008, several european agencies reported ml values for their stations. the values ranged from 4.6 to 5.8 even though the local corrections are supposed to ensure similar values for all stations (fig. 4). ideally the ml values estimated at seismographs map the energy released in different directions from the earthquake, but the uncertainty in the correction of damping is large, so the pattern is only determined with some uncertainty. at geus, a ml value of 4.8 was calculated, identical to the ml value reported by the usgs. university of bergen reported 4.7, university of helsinki 4.9 and helmholtz-zentrum potsdam deutsches geoforschungszentrum 5.4. the moment magnitude is based on the seismic moment of the earthquake, a physical parameter that is proportional to the area of the fault multiplied by its slip. the moment can be determined from seismological measurements, and in special cases from geodetic measurements. a large number of clear signals on seismographs surrounding the epicentre are needed to obtain a well constrained moment magnitude, using moment tensor inversion. the moment magnitude is determined by searching for the earthquake source parameters that produce synthetic seismograms that have the best possible fit to the observed seismograms. we have determined the moment tensor and a moment magnitude of 4.2 using the method by dreger (2003). university of uppsala, usgs and istituto nazionale di geofisica e vulcanologia reported a moment magnitude of 4.3. fault plane solutions determined in this preliminary study and by others all show a strike-slip movement with a normal component. the corresponding compressional stress field is orientated in a nnw–sse direction as shown by the focal mechanisms (fig. 4). 11 fig. 5. seismograms of the 16 december 2008 earthquake showing the vertical component of the broadband seismometers at copenhagen and bornholm, located 65 and 108 km from the epicentre, respectively. fig. 4. map showing magnitudes and focal mechanisms from various agencies. each circle shows the ml estimated from the data from each seismograph. ml values are from university of bergen, university of helsinki, geoforschungszentrum in postdam and geus. the 2008 epicentre is shown by a star. the focal mechanisms in the upper left corner were determined by this (preliminary) study, cmt (the global cmt project; ekström & nettles 2008) and ingv (the mednet network; morelli et al. 2000). 400 km 5.8 5.6 5.4 5.2 5.0 4.8 4.6 10°e0°10°w 64°n 60°n 56°n 52°n 48°n 20°e this study cmt ingv 20 21 copenhagen bornholm minutes rosa_2008:rosa-2008 01/07/09 15:47 side 11 tectonic setting the epicentre is located in the sorgenfrei–tornquist zone, which is part of a major transition zone in europe between (1) the old precambrian shield of fennoscandia and eastern europe and (2) the younger lithosphere of central europe (pharaoh 1999). the earthquake activity in the sor gen frei–tornquist zone is limited, with kattegat being the most active region. the most active area in denmark is from north-western jylland into skagerrak (gregersen et al. 1998). recent studies show that many of the earthquakes cannot be referred to known faults in the sor genfrei–tornquist zone (gregersen et al. 1996). this is the case even though the uncertainties in the earthquake locations are large. faults in the top pre-zechstein rocks have been mapped in the area where the earthquake occurred (fig. 1). however, the location uncertainty (6 km) means that one cannot link the 16 december 2008 earthquake to a known fault. this is a general problem as small earthquakes occur at depth where as faults are mostly mapped near the surface. a re-analysis of the data recorded by the danish seismological network shows that no foreshocks or aftershocks were observed in the months before or after the 16 december earthquake. seismograms of the earthquake are shown in fig. 5. concluding remarks to feel an earthquake in denmark is rare but not unlikely. earthquakes felt in denmark are generally small and felt in north-western jylland or northern sjælland (gregersen et al. 1998; larsen et al. 2008). large earthquakes with epicentres located outside denmark have also been felt in denmark, like the earthquakes in lissabon in 1755, oslo in 1904 and kaliningrad in 2004 (gregersen et al. 2007). the perceptibility area in denmark of the 16 december 2008 earthquake is comparable to the perceptibility area of the holbæk earthquake in 2001 (larsen et al. 2008), where the earthquake was felt mostly in northern sjælland. also, the area is not much dif ferent from that shaking in 1930 by a small earthquake in øresund. the shaking from the oslo earthquake in 1904 was felt in the north-eastern part of denmark. the kaliningrad earthquake in 2004 was felt by many in copenhagen and in northern sjælland, but only marginally in the rest of den mark. the coincidence of the perceptibility areas must be included in future evaluations of the earthquake hazard of denmark. denmark is located in an area of low earthquake hazard (giardini et al. 2003) and the danger of any damage or injury caused by an earthquake is insignificant compared to other hazards. acknowledgements university of bergen, university of helsinki, university of uppsala, deutsches geoforschungszentrum in postdam and usgs provided data for this study. references ekström, g. & nettles, m. 2008: the global cmt project, http:// www. globalcmt.org. dreger, d.s. 2003: tdmt_inv: time domain seismic moment tensor inversion. in lee, w.h.k., kanamori, h., jennings, p.c. & kisslinger, c. (eds): international handbook of earthquake and engineering seismology 81b, 1627 only. london: academic press. giardini, d., jiménez, m.-j. & grünthal, g. (eds) 2003: europeanmediterranean seismic hazard map. european seismological com mission. gregersen, s. 1992: crustal stress regime in fennoscandia from focal mechanisms. journal of geophysical research 97, 11,821–11,827. gregersen, s., korhonen, h. & husebye, e.s. 1991: fennoscandian dynamics: present-day earthquake activity. tectonophysics 189, 333–344. gregersen, s., leth, j., lind, g. & lykke-andersen, h. 1996: earthquake activity and its relationship with geologically recent motion in denmark. tectonophysics 257, 265–273. gregersen, s., hjelme, j. & hjortenberg, e. 1998: earthquakes in denmark. bulletin of the geological society of denmark 44, 115–127. gregersen, s., wiejacz, p., debski, w., domanski, b., assinovskaya, b., guterch, b., matyniemi, p., nikulin, v., pacesa, a. & puura, v. 2007: the exceptional earthquakes in kaliningrad district, russia on september 21, 2004. physics of the earth and planetary interiors 164, 63–74. grünthal, g., musson, r.m.w., schwarz, j. & stucchi, m. (eds) 1998: european macroseismic scale 15, 99 pp. luxembourg: cahiers du centre européen de géodynamique et de séismologie. hanks, t.c., & kanamori, h. 1979: a moment magnitude scale. journal of geophysical research 84(b5), 2348–2350. larsen, t.b., gregersen, s., voss, p.h., bidstrup, t., & orozova-bekkevold, v. 2008: the earthquake that shook central sjælland, denmark, november 6, 2001. bulletin of the geological society of denmark 56, 1–11. morelli, a., ekström, g., mazza, s., pondrelli, s., boschi, e. & dziewonski, a.m. 2000: surface-wave centroid moment tensors in the mediterranean region: the mednet-harvard project. orfeus electronic newsletter 2, p. 4. pharaoh, t. 1999: palaeozoic terranes and their lithospheric boundaries within the trans-european suture zone (tesz): a review. tectono physics 314, 17–41. richter c.f. 1935: an instrumental earthquake magnitude scale. bulletin of the seismological society of america 25, 1–32. scandinavian earthquake archive 2003: cd with scanned reports for icg project no. 3. icg 3-2003-3. authors’ addresses p.h.v., t.b.l. & s.g., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: pv@geus.dk l.o., british geological survey, murchison house, west mains road, edinburgh eh9 3la, uk. 12 rosa_2008:rosa-2008 01/07/09 15:47 side 12 geological survey of denmark and greenland bulletin 42, 2018, 149-168 149 late jurassic evolution of the jameson land basin, east greenland – implications of the blokelv-1 borehole morten bjerager, peter alsen, jørgen a. bojesen-koefoed, tove nielsen, stefan piasecki and anders pilgaard data from the recently drilled, fully cored blokelv-1 borehole and previous cored boreholes in the upper jurassic of jameson land, central east greenland, are integrated with published field studies to address the depositional evolution of the jameson land basin in the oxfordian–volgian. in jameson land, the succession represents a marine shelf-to-basin transect in a w–sw-dipping half-graben. laminated organic-rich mudstones were deposited in the central deep parts of the basin and grade up-slope into bioturbated sandy mudstones. extensive shallow marine – deltaic sand prograded from the western and northern basin margins and formed prominent sandy shelf-edge wedges. sand-rich density flows initiated by periodic collapse of the shelf edge deposited massive sand bodies on the slope and basin floor; these sands were prone to post-burial remobilisation to form injectite bodies. basin evolution was controlled both by relative sea-level changes, typically correlatable with regional and global sea-level curves, and by rift tectonics. during periods with high relative sea level, the organicrich muddy facies onlapped the sandy shelf environments; such periods of basinal expansion and onlap are recorded in the lower oxfordian (q. mariae chronozone), the middle–upper oxfordian (c. tenuiserratum – a. glosense chronozones) and uppermost oxfordian – upper kimmeridgian (a. regulare – a. autissiodorensis chronozones); the deepening, transgressive trend culminated in the mid-kimmeridgian (a. eudoxus chron). marked progradation of the sandy shelf and associated deposition of gravity-flow sands on the slope and basin floor occurred in the early oxfordian (c. cordatum chron), the middle oxfordian (c. densiplicatum chron), the late oxfordian (a. serratum chron) and the early volgian (p. elegans chron). the basin architecture reflects periodic differential subsidence on the wto sw-dipping fault block. the lower to middle oxfordian is highly condensed in the east (<10 m) and thickens markedly towards the west (>300 m), reflecting accumulation during rift/fault-controlled block rotation. the upper oxfordian – kimmeridgian, in contrast, shows a broadly symmetrical distribution and records uniform regional subsidence. keywords: hareelv formation, jameson land basin, blokelv-1, slope and basin floor deposition, relative sea level, sedimentary architecture ___________________________________________________________________________ geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: mbj@geus.dk the exposed jurassic succession in jameson land, and in eastern greenland in general, has a long history of geological research. in the last decades, it has acquired special relevance in a petroleum geological context as an analogue for similar basin settings offshore along the east greenland margin and on the conjugate norwegian continental shelf. the upper jurassic hareelv formation in jameson land represents a complex intercalation of organic-rich marine mudstones and potential reservoirquality gravity flow and injected sandstones in an appar© geus, 2018. geological survey of denmark and greenland bulletin 42, 149–168. available at: www.geus.dk/bulletin42 mailto:mbj@geus.dk http://www.geus.dk/bulletin42 ; ; ; ; ; ; bay fjelde blokel v savoia halvø hal l bredn ing kap leslie hartz fjeld charcot havn mudderbugten jameson land h urry in le t liverpool landmilne land a a' hors ens fjo rd carlsberg f jord na tho rst fj ord fle ming fjo rd katedralen fortet raukelv olympen parnas fossilbjerget mikael bjerg hareelv gåseelv ugleelv 22°w 72°n 71°n 72°n 71°30'n 70°30'n 23°w24°w25°w 22°w 25°w mbj artikel 2 fig 1 71°30'nmajor dyke/sill fault ice river quaternary hesteelv fm hartz fjeld fm paleogene basalts paleogene cenozoic raukelv fm hareelv fm, salix dal mb olympen fm fossilbjerget fm/pelion fm neill klinter gp kap stewart gp triassic charcot bugt fm kap leslie fm basement hareelv fm, sjællandselv mb hareelv fm, katedralen mb devonian carboniferous permian 20 km s c o r e s b y s u n d jyllandselv sjællandselv lollandselv lollandselv-2 lollandselv-1 falsterelv-1 jyllandselv-1 blokelv-1 sjællandselv-1-2 -3 falsterelv fig. 1. geological map of jameson land showing the locations of cored stratigraphic boreholes and selected outcrop localities. based on the digital greenland geological map at a scale of 1:500 000 and printed map series at a scale of 1:100 000; only named rivers are shown. a–a’ indicates the line of the cross-section in fig. 2. 151 ently non-predictive stratigraphical context (surlyk et al. 2007). upper jurassic sequence stratigraphy in the basin is thus based on the exposed succession in milne land (surlyk 1991; larsen et al. 2003), which is considered to represent a separate fault block within the jameson land basin (figs 1, 2). the presence of jurassic deposits in jameson land has been known since the early 1800s (madsen 1904; rosenkrantz 1929; aldinger 1935). large-scale geological mapping campaigns of the region in the mid-1900s (donovan 1957; haller 1971) were followed by systematic detailed mapping by the geological survey of greenland (ggu now geus) and the university of copenhagen, resulting in formal lithostratigraphic subdivision of the jurassic succession in jameson land (surlyk et al. 1973), later provisionally revised by surlyk (2003). shallow coring campaigns were conducted in 1982–1983 and 1993 by ggu, targeting upper jurassic potential source rocks (piasecki et al. 1996). the biostratigraphic subdivision of the succession has been predominantly based on ammonites collected over numerous field seasons, combined with palynomorph assemblages, mainly prepared from mudstone samples. a review of previous biostratigraphic data is included in the presentation of the biostratigraphic subdivision of the blokelv-1 core by alsen & piasecki (2018, this volume). modern sedimentological and stratigraphical studies on the upper jurassic based on field work were conducted over the last decades by larsen et al. (2003) in milne land and by surlyk (1987), surlyk & noe-nygaard (1991, 2000, 2001, 2003, 2005), larsen & surlyk (2003), bruhn & surlyk (2004) and surlyk et al. (2007) in jameson land. the main focus was on the coarse-grained, sand-rich units and their relationships to overlying and underlying mudstone units, whereas the mainly poorlyexposed mudstone successions attracted less attention. a comprehensive review and summary of the jurassic literature was presented by surlyk (2003). more recently, geochemical results from milne land were presented by strogen et al. (2005) and a biostratigraphic review on the jurassic was published by kelly et al. (2015). this paper provides new information on unweathered and well-preserved core material from the central part of jameson land represented by the recently released, 233 m long blokelv-1 core (bjerager et al. 2018a, this volume) and by cores, 30–100 m deep, from the earlier drilling campaigns (requejo et al. 1989; piasecki et al. 1996; pilgaard 2012). the thick and stratigraphically complete blokelv-1 core provides a stratigraphic link between the previously drilled cores through the upper jurassic succession and outcrops across jameson land and milne land (figs 3, 4). the stratigraphic and sedimentological analyses of the cores, integrated with published field studies, contribute to refinement of the understanding of the depositional evolution of a rift-controlled shelf– basin transect in the late jurassic. geological setting the jameson land basin is up to about 150 km wide and more than 200 km long and contains a nearly complete post-caledonian to lower cretaceous sedimentary succession (fig. 1). the jameson land basin thus contains a continental devonian basin succession up to 8 km thick, succeeded by n–s-trending and west-dipping carboniferous rift basins. early permian uplift resulted in the formation of a regional peneplain (kempter 1961), which fig. 2. geological profile (see location in fig. 1), based on larsen (1980), birkelund et al. (1984), larsen et al. 2003, surlyk (2003), guarnieri et al. (2017), reprocessed awi seismic lines in hall bredning (fechner 1994; geus unpublished data 2018), geological map 1:100 000 (geological survey of greenland), and the blokelv-1 core. cb: charcot bugt fm. fo: fossilbjerget fm. ha: hareelv fm. hf: hartz fjeld fm. kl: kap leslie fm. ks: kap stewart gp. nk: neill klinter gp. ol: olympen fm. pe: pelion fm. ra: raukelv fm. tr: triassic. kl cb hf basement ra ha ol fo pe nk ks tr pe fo basalt basement 1 al tit ud e (k m ) 0.5 –0.5 0 tr bay fjelde hartz fjeld ugleelv hurry inlet liverpool land sjællandselv jameson landmilne land charcot havn blokelv-1 -1-2-3 katedralen horsens fjordhall bredning mbj_artikel 2_fig2_sts 152152 was segmented during the development of late permian – early triassic rift basins. a regional structural change to ne–sw-oriented rift basins has been proposed for the early triassic (seidler et al. 2004; guarnieri et al. 2017). renewed rifting in jurassic times was oriented mainly n–s with a proposed major fault situated in the western part of the scoresby sund fjord (hall bredning) bounding a general west-dipping half-graben in jameson land (fig. 2). the milne land fault block was onlapped in the middle jurassic (larsen et al. 2003) whereas the permian peneplain on liverpool land was onlapped in the middle –late triassic and jurassic. the nature of the northern and southern limits of the jurassic basin is less clear. an inferred nw–se-trending fault in kong oscar fjord between jameson land and traill ø has been suggested to have influenced basin development (surlyk 1978), and may be a continuation of the major jan mayen fracture zone and lineament on the norwegian continental shelf (e.g. lundin & doré 1997; guarnieri et al. 2017). towards the south, the basin probably continues south of the scoresby sund fjord beneath the kilometrethick palaeogene basalt succession (larsen 1980; larsen & marcussen 1992; nøhr-hansen & piasecki 2002). in jurassic times, the basin evolved as an asymmetric sag with deposition of a >2 km thick succession in the deepest parts of the basin (surlyk 2003). a major n– s-oriented fault zone in hall bredning (inner scoresby sund) is suggested to have controlled the syn-rift depositional systems in the basin (figs 1, 2, 4). deposition in early–middle jurassic times took place in lacustrine and successively shallow marine shelf settings (dam & surlyk 1992, 1998). an internal slope developed in the oxfordian, resulting in differentiation into a shallow marine sandy shelf to the north, a slope and a southern deeperfig. 3. the middle–upper jurassic succession in the ugleelv valley in jameson land, viewed towards the east (liverpool land in the distance). shallow marine sandstones of the pelion formation (c. 100 m thick) overlie the middle juassic neill klinter group and are in turn overlain by grey silty mudstones of the fossilbjerget formation (c. 100 m thick). the latter is succeeded unconformably by black mudstones and intercalated yellow massive sandstones of the hareelv formation, 150 m thick. hareelv fm hareelv fm neill klinter gp fossilbjerget fm fossilbjerget fm pelion fm pelion fm gø06_606_mbj_blokelv mbj_artikel 2_fig3_sts 153 water basin setting (bruhn & surlyk 2004). the late jurassic evolution of the jameson land basin is considered analogous to basins offshore nw europe, as reflected in the stratigraphic correlation charts in surlyk (2003), surlyk & ineson (2003) and stoker et al. (2017). the late jurassic was thus characterised by an overall sea-level lowstand in the middle oxfordian, a highstand in the kimmeridgian and a lowstand in the volgian (surlyk 1991). a major unconformity showing valley incision marks the boundary to the marine lower cretaceous hesteelv formation, which forms the top of the mesozoic succession in jameson land (surlyk et al. 1973). palaeogene basaltic intrusions form prominent wnw–ese-trending dykes and sills in the basinal succession, and extrusive volcanics erosionally overlie the jurassic to the west in milne land. material and methods this study is based on sedimentological descriptions of seven fully cored shallow boreholes, 30–100 m deep, drilled in the 1980s and 1990s, and of the recent fully cored blokelv-1 borehole, 233 m deep (bjerager et al. 2018b, this volume). total gamma-ray logs are available for all except the sjællandselv cores. spectral gamma ray logs and a density log are available for the blokelv-1 core (bjerager et al. 2018b, this volume). the cores represent a nw–se-oriented transect from the shelf edge via the slope to the deep central part of the basin with the blokelv-1 core serving as a key correlation link, by virtue of its stratigraphic completeness (figs 4–6). sedimentological data, petrophysical log trends and ammonite and palynostratigraphical data from the cores are all integrated with published outcrop studies and geus in-house, unpublished field data from jameson fig. 4. stratigraphic scheme based on the geological timescale of gradstein et al. (2012,) showing a w–e cross-section from milne land to central jameson land, a s–n transect in jameson land and an inferred relative sea-level curve. as: astartedal mb. cb: charcot bugt fm. mu: mudderbugt mb. pe: pernaryggen mb. vi: visdal mb. modified from larsen et al. (2003) and surlyk (2003). c re ta ce ou s lo w er u pp er ju ra ss ic o xf or di an l l l l m u u u sy st em se rie s te th ys (s ta nd ar d) c hr on oz on es re lat ive se a le ve l bo re al su bs ta ge m ki m m er id gia n be rr ias ian ti th on ian ry az . vo lgi an age (ma) 150 160 145 155 stage gø02_02_175_mbj_blokelv.eps w s n bl ok el v1 bl ok el v1 jy lla nd se lv1 fa lst er el v1 lo lla nd se lve lv1 lo lla nd se lve lv2 sjæ lla nd se lve lv3 sjæ lla nd se lve lv2 sjæ lla nd se lve lv1 bay fjelde hartz fjeld kap leslie hall bredning raukelvgåseelv olympen e milne land jameson land jameson land r. c. a. s. a. g. c. t. c. d. c. c. q. m. p. b. p. e. p. w. p. s. p. p. krebsedal mb hartz fjeld fm c c c c sa pe gråkløft mb cardioceraskløft mb bays elv mb kap leslie fm al kosmocerasdal mb mu p. n. a. ro. a. m. a. eu. a. a. a. re. hennigryggen mb (lower part) shallow marine sandstone hiatus shelf transition silty–sandy mudstone and heterolith deep marine sandstone shelf–basin mudstone potential source rock condensed sectionmass-flow sandstone injected sandstone fault (inferred) prograding unit 140 km 120 km hareelv fm hareelv fm hareelv fm?hareelv fm? hades mbhades mb katedralen mb katedralen mb zeus mb olympen fmolympen fm sjællandselv mb raukelv fm raukelv fm vi cb as mbj_artikel 2_fig4_sts 154154 c fault condensedc gø01_06m_034_mbj deep-water marine sandstone bodies drowned sandy (relict) shelf (olympen fm, zeus mb)shallow marine sandstone offshore marine mudstone inferred coastline transition zone marine sandstone/ mudstone heterolith n 50 kmc d a b jameson land traill ø milne land mbj_artikel 2_fig5_sts fig. 5. late jurassic palaeogeography of the jameson land area, modified after surlyk (2003). a: early oxfordian q. mariae chronozone. b: early–middle oxfordian c. cordatum – c. densiplicatum chronozones. c: latest oxfordian – late kimmeridgian a. regulare – a. eudoxus chronozones. d: early volgian. 155 land and milne land and seismic data from hall bredning. this provides a robust and detailed stratigraphic framework at ammonite chronozone level (alsen & piasecki 2018, this volume) that forms the framework for the description of the late jurassic evolution of the basin described here. the cores together provide a record of the lower oxfordian – lower volgian. the three northern cores at falsterelv and lollandselv (ggu 303139–303141) document the lower–middle oxfordian, whereas the southern cores at jyllandselv (ggu 303142), blokelv (ggu 511101) and sjællandselv (ggu 303114–301116) span the middle oxfordian – lower volgian interval (figs 4, 6–9). regional depositional evolution the temporal evolution of the jameson land basin in the late jurassic as recorded by the integrated subsurface borehole data is based primarily on the biostratigraphic framework developed for these cored sections (piasecki et al. 1996; alsen & piasecki 2018, this volume). the combination of the deep-water nature of these sediments and the complex stratigraphic relationships developed in the extensively mobilised, intruded basinal successions precludes systematic sequence stratigraphic analysis, as has been successfully applied in the marginal facies in jameson land and milne land (larsen et al. 2003; larsen & surlyk 2003; surlyk & noe-nygaard 1991, 2000, 2005). although the following account is thus structured on the basis of the biostratigraphic subdivision into chronozones, recognition of sand-dominated and mud-dominated sedimentary units of comparable ages permits the correlation of certain sedimentary bodies between wells, despite the uncertainties associated with sediment mobilisation and intrusion (see figs 6, 8). lower oxfordian q. mariae and c. cordatum chronozones the succession referred to the q. mariae chronozone records a general regional sea-level rise in the basin (figs 4, 5a), such that dark grey mudstones of the kosmocerasdal member, up to 10 m thick, drape the basin margin in milne land (callomon & birkelund 1980; birkelund et al. 1984; larsen et al. 2003). in central jameson land, the chronozone comprises dark grey and black silty mudstones of the hades member, olympen formation (larsen & surlyk 2003; bruhn & surlyk 2004). the chronozone is also recognised in the basal part of the lollandselv-1 core, where it comprises bioturbated sandy mudstones (figs 6, 9), and at outcrop in the basal hareelv formation mudstones in the ugleelv (poulsen 1985) and gåseelv valleys. the c. cordatum chronozone is characterised by prominent shallow marine sand clinothems (c. 40 m thick) that prograde eastwards in milne land (unit cb4 of the visdal member, larsen et al. 2003). the correlative unit in jameson land is dominated by gravity-flow sand bodies that were transported from the shelf edge at the northern edge of the basin and down-slope towards the south and south-east. in the lollandselv-1 core, this chronozone is over 80 m thick (fig. 6; pilgaard 2012); individual sand units are 0.5 to 20 m thick, commonly show a downslope increase in thickness and are intercalated with bioturbated sandy mudstones. at olympen, the c. cordatum chronozone is not recognised biostratigraphically but this regressive phase may be represented by deeply incised gully-fill sediments at the top of the hades member (larsen & surlyk 2003). middle oxfordian c. densiplicatum and c. tenuiserratum chronozones the uppermost c. cordatum chronozone and the lower c. densiplicatum chronozone record a sea-level rise in the basin with deposition of offshore bioturbated sandy mudstones as recorded in the falsterelv and lollandselv cores (figs 4, 6) and traced to the southern part of traill ø (bruhn & surlyk 2004). in the proximal setting at olympen, sandy progradational deposits of the zeus member (olympen formation) up to 150 m thick, represent the deltaic shelf edge during the c. densiplicatum chron (figs 1, 4, 5b; larsen & surlyk 2003; bruhn & surlyk 2004). the upper c. densiplicatum chronozone defines a nw–se shelf-edge – slope – basin-floor transect: the zeus member shelf-edge facies correlate with massive gravity-flow sands (c. 10 m thick) in a slope setting at lollandselv and falsterelv passing into basin-floor turbidites and massive sands in the blokelv-1 core (>17 m thick; figs 6–8; bjerager et al. 2018b, this volume). at the western margin of the basin, in milne land, this interval is equivalent to a succession of shallow marine and fluvial sandstones and conglomerates (6 m thick) referred to the mudderbugt member (charcot bugt formation; larsen et al. 2003). the succeeding sea-level rise recorded in the c. tenuiserratum chronozone is represented by a succession of bioturbated sandy mudstones 156156 0 10 20 30 40 50 60 70 80 90 100 10 20 30 40 50 60 70 80 olympen fm nnw lo w er o xf or di an m id dl e o xf or di an hareelv fm lollandselv-2 (303140) lollandselv-1 (303139) falsterelv-1 (303141) c. tenuiserratum c. tenuiserratum c. densiplicatum c. densiplicatum a. glosense c. cordatum q. mariae 0 10 20 30 40 50 60 70 80 90 100 4.5 km 5.3 km 11.7 km 0 100 grtotal (api units) 0 100 grtotal (api units) 0 125 grtotal (api units) c. cordatum c. densiplicatum c. cordatum fig. 6. correlation panel (broadly nnw–sse) of the cored borehole sections in jameson land. grtotal (where available), lithological and biostratigraphic summary data are given for the individual cored sections; additional sedimentological data are presented in figs 7–9. 157 0 10 20 30 5 10 20 30 0 10 20 30 40 50 60 70 3 10 20 30 40 sjællandselv-2 (303115) sjællandselv-1 (303114) sjællandselv-3 (303116) jyllandselv-1 (303142) hareelv fm (katedralen mb) hareelv fm (sjællandselv mb) sse u pp er o xf or di an lo w er k im m er id gia n u pp er k im m er id gia n m id dl e o xf or di an lo w er v ol gia n blokelv-1 (511101) 10 h a re e lv f m 20 30 40 50 60 70 80 90 100 110 p. elegans a. autissiodorensis a. eudoxus a. mutabilis p. baylei – r. cymodoce r. cymodoce p. baylei p. baylei a. regulare – a. rosenkrantzi a. regulare – a. rosenkrantzi a. glosense – a. serratum c. tenuiserratum 3.9 km 11.3 km c. densiplicatum 120 130 140 150 160 170 180 190 200 210 220 230 233.8 r. cymodoce mudstone, laminated (basin floor) sandy, silty mudstone, bioturbated (slope) heterolithic sandstone/mudstone (slope – basin floor) sandstone, homogeneous (slope – basin floor) sandstone (injectite) lithology and depositional environment igneous intrusion belemnite ammonite bivalve brachiopod correlation line biostratigraphic correlation line 0 150 grtotal (api units) grtotal (api units) 0 100200 158158 chrono. standard chronozones gø04_02m_453_01_mbj_blokelv_reg mbj_artikel 2_fig7_sts u pp er ju ra ss ic ki m m er id gia n up pe r up pe r lo w er lo w er o xf or di an p. elegans a. autissiodorensis a. mutabilis p. baylei – r. cymodoce p. baylei a. regulare – a. rosenkrantzi a. glosense – a. serratum c. tenuiserratum c. densiplicatum volgian ? a. eudoxus r. cymodoce 0 0 5 10 15100 200 grtotal (api units) blokelv-1 (511101) u (ppm) 20 depth (m) 10 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 210 220 230 10 depth (m) 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 210 220 230 233.8 be be mudstone heterolith (mudstone/sandstone) sandstone sandstone, remobilised (intruded) lithology igneous intrusion be ripple cross-lamination wavy bedding slump sandstone intrusion small mudstone intraclasts bentonite large mudstone clast coalified wood belemnite ammonite bivalve brachiopod structures, biota parallel lamination/ bedding diffuse stratification clay si sand pebbl. fig. 7. sedimentological log of the blokelv-1 core showing the biostratigraphical subdivision together with total and spectral (uranium) gamma-ray logs. for additional biostratigraphical and sedimentological data, see alsen & piasecki (2018, this volume) and bjerager et al. (2018b, this volume). 159 (c. 20 m thick) in the slope setting at lollandselv and laminated basinal mudstones (c. 9 m thick) at the blokelv-1 location (figs 6, 8–10). upper oxfordian a. glosense, a. serratum, a. regulare and a. rosenkrantzi chronozones the a. glosense chronozone records a further rise in sea level, the zone being represented by black, laminated basin-floor mudstones (katedralen member, hareelv formation) in the blokelv-1 and falsterelv-1 cores. similar facies crop out at ugleelv (surlyk et al. 1973) and gåseelv (fig. 1), and offshore bioturbated sandy siltstones of the kosmocerasdal member comprise the chronozone in milne land. a prominent relative sea-level fall is represented in the succeeding a. serratum chronozone, as testified by shallow marine sandstones (0–70 m thick) of the aldinger elv member (‘pecten sandstone’) in milne land (callomon & birkelund 1980; larsen et al. 2003), and by basin-floor, gravity-flow sandstones at blokelv (figs 4, 6, 7). the combined thickness of the a. glosense – a. serratum zones is c. 48 m in blokelv-1, 40 m at ugleelv and 10–80 m in milne land. the upper part of the a. serratum chronozone in the blokelv-1 core displays a marked increase in the general spectral gamma-ray uranium signal in the mudstones. this culminates in the lower part of the a. regulare chronozone, and possibly, also in the lower part of the a. rosenkrantzi chronozone, reflecting increasingly poorly oxygenated conditions on the sea floor (bjerager et al. 2018b, this volume; fig. 7). sedimentation was dominated by the accumulation of black laminated muds and intercalated homogenous gravity-flow sands that were largely remobilised post-depositionally and injected into the muds, a feature characteristic of the katedralen member of the hareelv formation (surlyk & noe-nygaard 2001; surlyk et al. 2007). coeval deposition in milne land is represented by offshore sandy mudstones (bays elv member) although these sediments were deposited in a considerably shallower water setting and have no source-rock potential (strogen et al. 2005), in contrast to the katedralen member in the blokelv and jyllandselv cores (bojesen-koefoed et al. 2018, this volume). the jyllandselv-1 core contains markedly less gravity flows and sandstone injectites compared with the corresponding interval in the blokelv-1 core (fig. 6), hinting at the considerable lateral variability of this unit, a feature that is also evident in exposures (fig. 11). the combined thickness of the a. regulare – a. rosenkrantzi chronozones is 15–40 m thick in milne land (callomon & birkelund 1980), 39 m in blokelv-1 and 23 m in jyllandselv-1. the upper oxfordian – volgian is absent north of falsterelv in jameson land (fig. 1). it thus remains open to speculation whether the sandy shelf continued to prograde southwards, or was aggradational or even retrogradational in the late oxfordian (fig. 4). the former middle oxfordian sandy shelf area may have formed a starved, bypass shelf in the late oxfordian with the depocentre shifting to the deeper part of the basin, farther south. regardless of the nature of the sandy shelf development, the oxfordian provided a source area for the massive gravity-flow sands deposited in basinal areas due to collapse of the sandy shelf margin. this situation lasted well into the kimmeridgian (figs 4, 5c). lower kimmerigian p. baylei and r. cymodoce chronozones only the deeper parts of the basin in jameson land have deposits preserved from this interval, i.e. from jyllandselv and southwards (figs 1, 4, 6). the deposits are similar to those of the underlying chronozones, being dominated by basinal mudstones in the blokelv-1 and jyllandselv cores and represented by intervals 33 m thick and 34 m thick, respectively (fig. 6). an overall increase succeeded by a decrease in the gamma-ray trend is recorded in the mudstone succession (fig. 7). as in the late oxfordian, the shelf area to the north probably represented a starved shelf with condensed deposition (fig. 5c). the sandy shelf edge became unstable and was subject to collapse during minor sea-level falls or during minor tectonic events, thereby serving as a source for gravity flows that transported sand-rich sediment downslope to the basin floor (surlyk et al. 2007). in milne land, the offshore mudstones of the bays elv member are overlain by sandy mudstones of the cardioceras kløft member at the boundary between the p. baylei and r. cymodoce chronozones suggesting a relative shallowing of the basin at this time (callomon & birkelund 1980); the combined thickness is 30–40 m in this area. upper kimmeridgian a. mutabilis, a. eudoxus and a. autissiodorensis chronozones this interval is recorded in the blokelv-1 core (78 m thick) and in the lower levels of the sjællandselv-3 core 160160 (10 m observed) and is assigned to the katedralen member (fig. 6). it comprises roughly equal proportions of interbedded black, organic-rich, laminated mudstones yielding very high gamma-ray values and massive gravityflow sandstones and injectites. heterolithic mudstones and sandstones dominate the upper 13 m of the unit (fig. 6). the equivalent interval in milne land is 80–110 m thick where the boundary between the sandy mudstones of the cardioceras kløft member and the black mudstones of the gråkløft member (potential source rocks) corresponds to the boundary between the a. mutabilis and a. eudoxus chronozones (birkelund et al. 1984; strogen et al. 2005). according to surlyk (2003), a peak sea-level highstand prevailed in the region during the a. eudoxus chron, an event recorded widely in the north atlantic region (surlyk 2003; haq 2017). in the blokelv-1 core, the uranium gamma-ray signal shows the highest consistent values in mudstones around the a. mutabilis – a. eudoxus chronozone boundary and in the lower part of the a. eudoxus chronozone. mudstones in the upper levels of the a. eudoxus chronozone show an upward-decreasing uranium gamma-ray trend and are progressively interbedded with prominent units of gravity-flow sandstones; these features may reflect a relative sea-level fall (figs 4, 6, 7). mudstones in the uppermost kimmeridgian a. autissiodorensis chronozone form five discrete packets, typically c. 5 m thick, that show upward-increasing gamma-ray trends followed by thinner decreasing trends; successive units show decreasing peak uranium levels (fig. 7). these cycles may record climate-induced variations in the production of organic algae (e.g. tyson 1996) or alternatively represent the autocyclic waxing and waning of detrital muddy sediment input. the mudstone packets are capped by gravity-flow sandstone units that were to some extent remobilised and injected after burial. lower volgian p. elegans chronozone the lower volgian is present in the blokelv-1 and sjællandselv cores and documents a basinal deepening trend from north to south. medium-grained massive sandstones of the sjællandselv member characterise the base of the p. elegans chronozone in the blokelv-1 core; the grain size of these sandstones is consistently coarser than that of the gravity-flow sandstones of the underlying katedralen member (fig. 7). outcrops adjacent to the blokelv-1 drilling location indicate that the sandstone units are laterally continuous over several kilometres and are interpreted to represent base-of-slope mass flows derived from the collapse of sandy high-angle clinothems in the shallow marine raukelv formation (surlyk 2003; surlyk & noe-nygaard 2005). the sjællandselv member sandstones do not show injection features in the vicinity of the blokelv-1 drill site. the lower volgian in the sjællandselv cores, however, comprises basinal mudstones and remobilised sandstones of the katedralen member, reflecting a more basinal location relative to the blokelv-1 core. in milne land, the p. elegans chronozone consists of deep-water, organic-rich mudstones of the gråkløft member. lower–middle volgian deposition of basinal organic-rich mudstone continued in the southernmost part of the jameson land basin contemporaneously with south-eastward progradation of coarse sandy clinoform beds of the raukelv formation from the early volgian p. pectinatus chron to the middle volgian. four major prograding units are recorded in the southern part of jameson land where they represent a major basin-filling phase (surlyk & noe-nygaard 1991; surlyk & noe-nygaard 2005). in the western, proximal part of the area (milne land), coeval shallow marine deposits are represented by the krebsedal, pernaryggen and astartedal members (kap leslie formation) and the lower part of the hennigryggen member (hartz fjeld formation) (birkelund et al. 1984). both the raukelv formation in jameson land and the lower hartz fjeld formation in milne land are capped by a major erosional unconformity that formed during an important sea-level fall in the latest volgian – earliest ryazanian (surlyk 1991). discussion depositional sequences and relative sea level sequence stratigraphic and tectonostratigraphic subdivisions of the jurassic – lowermost cretaceous succession based on large-scale depositional systems in the jameson land basin were presented by surlyk (1991, 2003). detailed sequence stratigraphic models based on the bajocian–oxfordian interval in milne land and the callovian–oxfordian and early–middle volgian of jameson land were presented by larsen et al. (2003), larsen & surlyk (2003) and surlyk & noe-nygaard (1991, 2000, 161 0 10 20 30 40 50 60 70 80 90 100 10 20 30 40 50 60 70 80 mud sand gr mud sand gr mud sand gr lollandselv-2 (303140) slope basin falsterelv-1 (303141) blokelv-1 (511101) a. glosense – a. serratum c. tenuiserratum c. tenuiserratum 8,5 km 14,9 km c. densiplicatum c. densiplicatum a. glosense c. cordatum 190 200 210 220 230 233.8 td c. cordatum 0 50 100 grtotal (api units) 0 50 100 grtotal (api units) 0 100 200 grtotal (api units) mbj_artikel 2_fig8_sts ripple cross-lamination wavy bedding slump sandstone intrusion small mudstone intraclasts bioturbation large mudstone clast coalified wood belemnite ammonite bivalve brachiopod shells structures, biota parallel lamination/ bedding diffuse stratification mudstone, laminated (basin floor) sandy, silty mudstone, bioturbated (slope) heterolithic sandstone/mudstone (slope – basin floor) sandstone, homogeneous (slope – basin floor) sandstone (injectite) lithology and depositional environment fig. 8. sedimentological and gamma-ray logs of the cored sections lollandselv-2, falsterelv-1 and the lower part of blokelv-1 showing biostratigraphic and lithostratigraphic correlations. 162162 2005), respectively. the succession in milne land was deposited in a proximal setting at the margin of the jameson land basin, and furthermore accumulated on a different fault block than the central parts of the jameson land basin (fig. 2). the cored sections in jameson land, however, facilitate detailed comparison of the depositional evolution of the proximal milne land and jameson land sediments with the deeper-water basinal facies of jameson land in the early oxfordian – early volgian period. regressive phases in the shelf-margin successions correlate broadly with basinal intervals dominated by gravity-flow sandstone units, as constrained by the 0 10 20 30 5 10 20 30 0 10 20 30 40 50 60 70 0 10 20 30 40 50 60 70 80 90 100 3 10 20 30 40 0 50 100 0 50 100 150 grtotal (api units) grtotal (api units) sjællandselv-2 (303115) sjællandselv-1 (303114) sjællandselv-3 (303116) jyllandselv-1 (303142) lollandselv-1 (303139) mbj_artikel 2_fig9_sts mud sand gr mud sand gr mud sand gr mud sand gr mud sand gr ripple cross-lamination wavy bedding slump sandstone intrusion small mudstone intraclasts bioturbation large mudstone clast coalified wood belemnite ammonite bivalve structures, biota parallel lamination/ bedding diffuse stratification mudstone, laminated (basin floor) sandy, silty mudstone, bioturbated (slope) heterolithic sandstone/mudstone (slope – basin floor) sandstone, homogeneous (slope – basin floor) sandstone (injectite) lithology and depositional environment shells fig. 9. sedimentological logs of the cored sections lollandselv-1, jyllandselv-1, sjællandselv-1, -2 and -3. 163 biostratigraphic results from the underlying and overlying mudstone units (fig. 7). the potential complications in correlation associated with remobilised, intrusive sand bodies are acknowledged but these broad correlations are supported by gamma-ray trends in the enveloping mudstones, as discussed below. the resulting relative sea-level curve (fig. 4) compares well with regional and global sea-level curves in the literature (e.g. hesselbo 2008; haq 2017). a marked sea-level rise characterised the early oxfordian (c. mariae chron) as reflected by regional deposition of offshore muds (fig. 5a). it was followed by a fall in relative sea level, recorded by a progradational unit of shallow marine sandstones in the lower–middle oxfordian c. cordatum and c. densiplicatum chronozones in milne land (larsen et al. 2003). the unit correlates with regressive shelf-edge deltaic sandstones of the zeus member in the olympen area and with interbedded bioturbated mudstones and gravity-flow sandstones in a slope setting in central jameson land (figs 4–9). a subsequent rise in relative sea level and regional drowning of the sandy shelf resulted in deposition of muds during the c. tenuiserratum and a. glosense chrons. a marked fall in relative sea level is inferred in the a. serratum chron, testified by the occurrence of shallow marine sandstones of the aldinger elv member in the eastern part of milne land (fürsich & heinberg 1983). this was contemporaneous with the deposition of prominent basin-floor, gravity-flow sandstones in jameson land that were remobilised and injected into encasing mudstones after burial. regional deposition of muds characterised the latest oxfordian (a. regulare chron), signalling renewed drowning of the basin that culminated with a peak sealevel highstand (maximum flooding interval) in the a. eudoxus – a. autissiodorensis chrons; maximum uranium gamma-ray levels are recorded in mudstones of this age in the jameson land basin (figs 4, 5c, 7). several units of mass-flow sandstones are intercalated in this succession; they commonly succeed mudstone intervals that show upward-decreasing uranium gamma-ray values and a concomitant increasing proportion of thin sand stringers, suggesting an inverse relationship between siliciclastic influx and organic (uranium) content. a reverse pattern – i.e. an upward-increasing gamma-ray trend – is commonly recorded above the massive mass-flow sandstone units. these patterns are evident despite the fact that most of the sandstone units are interpreted as having been injected post burial; this suggests that much of the sand was displaced only locally during post-depositional remobilisation and injection. the lower volgian p. elegans chronozone records the onset of a pronounced regressive phase during which extensive gravity-flow sand bodies in jameson land (now referred to the sjællandselv member) were deposited in front of the southward-prograding, sub-wave base, sandy clinothems of the raukelv formation (fig 5d; surlyk & noe-nygaard 2005). coeval deposition in milne land is represented by offshore sandy–silty mudstones of the krebsedal member. the pronounced progradation of the lower–middle volgian sandstones of the raukelv formation records a marked, stepwise sea-level fall in the region that was partly contemporaneous with deposition of the shallow marine sandstones of the pernaryggen member (kap leslie formation) and the lower part of the hartz fjeld formation in milne land (birkelund et al. 1984). accumulation rates and depositional architecture the overall basin morphology involved a west-dipping half-graben in jameson land bounded by a roughly n— s-trending syn-rift fault in hall bredning against the eroded and flooded basement block in milne land. this configuration is interpreted to have exerted a major influence on the depositional units and sediment geometries in the basin in late jurassic times (figs 1, 2, 4, 5; surlyk 1991). the biostratigraphic subdivision at chronozone level obtained from the core material and previous field studies can be integrated with the geological timescale of gradstein et al. (2012). conservative estimates of the post-compaction average sedimentation rates can thus be calculated for various basin settings and depositional units. it is acknowledged that individual calculated average accumulation rates at chronozone/substage level are associated with considerable uncertainties and absolute values should be treated with caution. it is considered instructive, however, to consider the relative variation in these values, both stratigraphically and geographically, as an additional guide to the factors controlling sedimentary processes and architecture in a rift-influenced basin. a conservative estimate of the average sedimentation rates of the combined post-compaction mudstone and sandstone can be calculated for the c. tenuiserratum to a. autissiodorensis chronozones in the blokelv-1 core. this stratigraphic interval includes no recognisable hiati and is considered to represent a complete and continuous sedimentary record from 159.9 to 152.1 ma (middle oxfordian to kimmeridgian–volgian boundary). this succession has a thickness of about 190 m in the blokelv-1 164164 toptop toptop 10 c m gø06_607_mbj_blokelv basebase basebase mbj_artikel 2_fig10_sts a b fig. 10. core photographs from the middle oxfordian c. tenuiserratum chronozone. a: slope facies from the lollandselv-2 borehole (26.75–22.75 m) comprising laminated and massive gravity-flow sandstones overlain by mid–dark grey biomottled silty and sandy mudstone. b: basin-floor facies from the blokelv-1 borehole (222.07– 218.25 m) consisting of massive, structureless gravity-flow sandstones overlain by laminated very dark grey to black mudstone. 165 core, yielding an overall average post-compaction accumulation rate of 2.5 cm/kyr of combined mudstone and sandstone deposition in the central part of the basin. the accumulation rate of the chronozones can also be estimated individually. the lower oxfordian q. mariae chronozone represents highly condensed basinal mud sedimentation (0–10 m) in the eastern part of the jameson land basin (poulsen 1985; bruhn & surlyk 2004), with an average sedimentation rate of <0.5 cm/kyr. the interval thickens to >50 m towards the west (larsen & surlyk 2003). the succeeding lower–middle oxfordian interval (c. cordatum – c. tenuiserratum chronozones) also shows a highly asymmetric depositional architecture. in the eastern part of jameson land, the mudstone-dominated basinal succession is 40 m thick with an average sedimentation rate of 2.5 cm/kyr. in contrast, in the northern part of jameson land (olympen), the coeval succession comprises up to 150 m of sandstone-dominated shelfslope deposits (bruhn & surlyk 2004). in the central and western part of the jameson land basin, the slope to basinal succession may be up to 200 m thick, estimated from seismic data, representing an average sedimentation rate of up to 15 cm/kyr (figs 2, 6). the wedge-shaped architecture of the lower–middle oxfordian succession in jameson land, thickening westwards, is suggestive of a rift-related control on sedimentation in a west-dipping half-graben setting with the controlling fault being situated in the offshore region of hall bredning (figs 2, 4, 5). in milne land, in a separate structural domain (fig. 2), this stratigraphic interval comprises a proximal to distal sedimentary package, 45–60 m thick. the lower part of the upper oxfordian (a. glosense – a. serratum chronozones) comprises a uniform mudstone-dominated basinal succession, 40–48 m thick, in the central and eastern part of jameson land representing average sedimentation rates of 6–7 cm/kyr. in milne land, this interval shows a wedge-shaped sedimentary geometry, 5–80 m thick. the uniform geometry of the uppermost oxfordian – kimmeridgian of the jameson land basin suggests that regional subsidence prevailed at this time. the upper oxfordian (a. regulare – a. rosenkrantzi chronozones) is dominated by a mudstone succession 15–40 m thick (thinnest at basin margins), representing average sedimentation of rates of 1–2 cm/kyr. the mudstonedominated basinal deposits of the lower and upper kimmeridgian show consistent thicknesses throughout the basin – 33–40 m and 80–110 m thick, respectively, giving average sedimentation rates of 1–1.5 cm/kyr and 3–4 cm/kyr. fig. 11. black mudstone and irregular, injected sandstone bodies (boundaries outlined) of the katedralen member (hareelv formation) in the ugleelv valley. vertical section illustrated is about 10 m thick. mbj_artikel 2_fig11_sts 166166 the jameson land basin experienced a remarkable shelf-edge evolution during the late jurassic, over a time span of c. 15 million years. in the middle oxfordian, the shelf edge was positioned in the olympen area as defined by the deltaic sandstone shelf-edge deposits of the zeus member (fig. 4; larsen & surlyk 2003; bruhn & surlyk 2004). by the volgian, the shelf edge had prograded over 100 km towards the south where it is recorded by the raukelv formation (fig. 5d; surlyk & noe-nygaard 2005). the position of the shelf-edge during deposition of much of the hareelv formation is uncertain, however, because correlative shallow marine sediments have been removed by modern erosion. previous palaeogeographic reconstructions avoided placing a sandy shelf-edge across jameson land in late oxfordian and kimmeridgian times despite the common occurrences of mass-flow sandstones in the basinal areas (figs 4, 7; surlyk 2003). it is proposed here, however, that the presence of mass-flow sandstones in basinal settings testifies to the existence of a sand-rich shelf edge, either a degrading relict middle oxfordian feature or a continuously active sand-prone shelf edge that was supplied with sand during the kimmeridgian (fig. 5c). the volgian saw a marked southward shift of the shelf edge with deposition of the shallow marine sandstones of the raukelv formation and the correlative basin-floor sandstones of the sjællandselv member (surlyk & noe-nygaard 2005). conclusions the exposed late jurassic record in the jameson land basin in east greenland provides the main elements of a syn-rift shelf-to-basin transect. deep-water basinal sediments consist of black, laminated, organic-rich mudstones intercalated with gravity-flow sandstones that commonly experienced post-burial remobilisation and injection; the shelf and shelf edge are characterised by sandy deltaic deposits. integration of these outcrop data with subsurface data derived from fully cored shallow boreholes yields a better understanding of the shelf-tobasin transition, particularly of the intermediate slope setting characterised by bioturbated sandy mudstones and gully-fill, gravity-flow sandstones. the 233.8 m thick blokelv-1 core represents a continuous middle oxfordian to early volgian stratigraphic record in a deep-water setting. this section provides the detailed stratigraphic control to relate the stratigraphic distribution of prominent sandstone units in the jameson land basin. despite remobilisation and injection, the development of a robust biostratigraphy and log stratigraphy supports the correlation of amalgamated sand bodies in the basinal succession with known pulses of shallow marine sand progradation in milne land and western and northern jameson land. major phases of sandy shelf progradation are recorded in the middle oxfordian c. densiplicatum chronozone, the upper oxfordian a. serratum chronozone, and in the volgian. the sand-prone shelf-edge in the jameson land basin shows a basin axial migration of >100 km towards the south between the middle oxfordian and the volgian, a time span of 15 million years. inferred deepening and transgressive phases are recorded in the lower oxfordian q. mariae chronozone, the middle–upper oxfordian c. tenuiserratum – a. glosense chronozones, the uppermost oxfordian a. regulare chronozone, and culminating in the upper kimmeridgian a. eudoxus – a. autissiodorensis chronozones where spectral log data suggest peak basin deepening and expansion (a key ‘maximum flooding interval’). the integrated subsurface and outcrop dataset documents a late jurassic history of relative sea-level variation that compares favourably with regional and global sea-level curves and confirms the jameson land basin as an important reference for late jurassic evolution of the north atlantic region. acknowledgements rasmus rasmussen is thanked for reprocessing of awi reflection seismic profiles from hall bredning. graphic work has been undertaken by jette halskov and stefan sølberg. we thank the referees rikke bruhn and finn surlyk for constructive and pertinent comments that contributed significantly to the revised paper. references aldinger, h. 1935: geologische beobachtungen im oberen jura des scoresbysundes (ostgrönland). meddelelser om grønland 99, 128 pp. alsen, p. & piasecki, s. 2018: biostratigraphy of the hareelv formation (upper jurassic) in the blokelv-1 core, jameson land, central east greenland. in: ineson, j. & bojesen-koefoed, j.a. 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(eds): petroleum geology of the upper jurassic – lower cretaceous of east and north-east greenland: blokelv-1 borehole, jameson land basin. geological survey of denmark and greenland bulletin 42, 39–64 (this volume). bjerager, m., bojesen-koefoed, j. & piasecki, s. 2018a: the upper jurassic blokelv-1 cored borehole in jameson land, east greenland – an introduction. in: ineson, j. & bojesen-koefoed, j.a. (eds): petroleum geology of the upper jurassic – lower cretaceous of east and north-east greenland: blokelv-1 borehole, jameson land basin. geological survey of denmark and greenland bulletin 42, 7–14 (this volume). bojesen-koefoed, j.a., peter nytoft, h.p., petersen, h.i., piasecki, s. & pilgaard, a. 2018: petroleum potential of the upper jurassic hareelv formation, jameson land, east greenland. in: ineson, j. & bojesen-koefoed, j.a. 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(eds): magmatism and the causes of continental break-up, geological society special publication 68, 365–186. larsen, m. & surlyk, f. 2003: shelf-edge delta and slope deposition in the upper callovian – middle oxfordian olympen formation, east greenland. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 931–948. larsen, m., piasecki, s. & surlyk, f. 2003: stratigraphy and sedimentology of a basement-onlapping shallow marine sandstone succession, the charcot bugt formation, middle–upper jurassic, east greenland. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 893–930. lundin, e.r. & doré, a.g. 1997: a tectonic model for the norwegian passive margin with implications for the ne atlantic: early cretaceous to break-up. journal of the geological society, london 154, 545–550. madsen, v. 1904: on jurassic fossils from east greenland. meddelelser om grønland 29(6), 157–211. nøhr-hansen, h. & piasecki, s. 2002: paleocene sub-basaltic sediments on savoia halvø, east greenland. geology of greenland survey bulletin 191, 111–116. pilgaard, a. 2012: unikke sandede tyngdestrømsaflejringer og injektitter i øvre jura, jameson land, østgrønland, 98 pp. unpublished thesis, københavns universitet, danmark. piasecki, s., dalhoff, f. & boserup, j. 1996: shallow boreholes from the jurassic of jameson land: analytical data. danmarks og grønlands geologiske undersøgelse rapport 30, appendix 9. poulsen, n.i. 1985: dinocyststratigrafien i den nedre del af hareelv formationen (øvre jura), jameson land, østgrønland. dansk geologisk forening årsskrift for 1984, 133–137. requejo, a.g., hollywood, j. & halpern, h.i. 1989: recognition and source correlation of migrated hydrocarbons in upper jurassic hareelv formation, jameson land, east greenland. aapg bulletin 73, 1065–1088. rosenkrantz, a. 1929: preliminary account of the geology of the scoresby sund district. meddelelser om grønland 73(2), 135–154. seidler, l., steel, r.j., stemmerik, l. & surlyk, f. 2004: north atlantic http://dx.doi.org/10.1130/gsatg359a.1 http://dx.doi.org/10.1130/gsatg359a.1 168168 marine rifting in the early triassic: new evidence from east greenland. journal of the geological society, london 161, 583–592. stoker, m. s., stewart, m. a., shannon, p. m. bjerager, m., nielsen, t., blischke, a., hjelstuen, b. o., gaina, c., mcdermott, k. & ólavsdóttir, j. 2017: an overview of the late paleozoic–mesozoic stratigraphy of the ne atlantic region. geological society special publication 447, http://dx.doi.org/10.1144/sp447.2. strogen, d.p., burwood, r. & whitham, a.g. 2005: sedimentology and geochemistry of late jurassic organic-rich shelfal mudstones from east greenland: regional and stratigraphic variations in source-rock quality. in: doré, a.g. & vining, b. (eds): petroleum geology: north-west europe and global perspectives – proceedings of the 6th petroleum geology conference, 923–932. geological society, london. surlyk, f. 1978: jurassic basin evolution of east greenland: nature 274, 130–133. surlyk, f. 1987: slope and deep shelf gully sandstones, upper jurassic, east greenland. aapg bulletin 71, 464–475. surlyk, f. 1991: sequence stratigraphy of the jurassic – lowermost cretaceous of east greenland. the aapg bulletin 75 (9), 1468– 1488. surlyk, f. 2003: the jurassic of east greenland: thermal subsidence, onset and culmination of rifting. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 659–722. surlyk, f. & ineson, j.r. 2003: the jurassic of denmark and greenland: key elements in the reconstruction of the north atlantic jurassic rift system. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 9–20. surlyk, f. & noe-nygaard, n. 1991: sand bank and dune facies architecture of a wide intracratonic seaway: late jurassic – early cretaceous raukelv formation, jameson land, east greenland. in: miall, a.d. & tyler, n. (eds): the three-dimensional facies architecture of terrigenous clastic sediments and its implications for hydrocarbon discovery and recovery. sepm (society for sedimentary geology). concepts in sedimentology and paleontology 3, 261–276. surlyk, f. & noe-nygaard, n. 2000. jurassic sequence stratigraphy of east greenland. in: hall, r.l. and smith, p.l. (eds): proceedings of the 5th international symposium on the jurassic system. georesearch forum 6, 357–366. trans tech publications ltd. surlyk, f. & noe-nygaard, n. 2001: sand remobilisation and intrusion in the upper jurassic hareelv formation of east greenland. bulletin of the geological society of denmark 48, 211–230. surlyk, f. & noe-nygaard, n. 2003: a giant sand injection complex: the upper jurassic hareelv formation of east greenland. geologia croatica 56, 69–81. surlyk, f. & noe-nygaard, n. 2005: a forced regressive shelf-margin wedge formed by transition-slope progradation: lowermost cretaceous rauk plateau member, jameson land, east greenland. bulletin of the geological society of denmark 52, 227–243. surlyk, f., callomon, j.h., bromley, r.g. & birkelund, t. 1973: stratigraphy of the lower jurassic – lower cretaceous sediments of jameson land and scoresby land, east greenland. grønlands geologiske undersøgelse bulletin 105, 76 pp. surlyk, f., gjelberg, j. & noe-nygaard, n. 2007: the upper jurassic hareelv formation of east greenland: a giant sedimentary injection complex. in: hurst, a. and cartwright, j. (eds): sand injectites: implications for hydrocarbon exploration and production. aapg memoir 87, 141–149. tyson, r.v. 1996: sequence-stratigraphic interpretation of organic facies variations in marine siliciclastic systems: general principles and application to onshore kimmeridge clay formation, uk. in: hesselbo, s.p. & parkinson d.n. (eds): sequence in stratigraphy in british geology, geological society special publication 103, 75–96. _________________________________________________________________________________________ manuscript received 16 december 2015; revision accepted 8 february 2018 http://dx.doi.org/10.1144/sp447.2 geological survey of denmark and greenland bulletin 15, 2008, 33-36 tunnel valleys eroded by subglacial meltwater underneath the late weichselian ice sheet are a common feature in the danish landscape (ussing 1907; smed 1998). they occur as undulating elongate depressions with hollows and thresholds and without continuously descending floors. the valleys rise tens of metres before terminating in large outwash fans, primarily along the main stationary line in jylland, but also along younger ice-margin lines formed shortly after the last glacial maximum. the meltwater was driven by hydrostatic pressure gradients below the glacier towards its margin leading to subglacial erosional features, partly in the form of valleys. the term ‘tunnel valley’ was first used by madsen (1921), who referred to tunnel-like structures below glaciers that were expected to have carried the meltwater. worldwide, this term is used for subglacially eroded valleys; however, other terms such as ‘tunnel channel’ and ‘incision’ are also widely used for such valleys. there is general consensus that subglacial meltwater is the primary causative agent that has eroded the tunnel valleys (o’cofaigh 1996; huuse & lykke-andersen 2000; jørgensen & sandersen 2006). the subglacial origin is indicated by: (1) abrupt terminations at former ice margins and the association with the large outwash plains, (2) irregular longitudinal profiles, (3) the occurrence of small channels and eskers in the valleys, and (4) the non-meandering and non-dendritic appearance of the relatively straight-segmented valleys. the exact mode of meltwater erosion remains, however, poorly understood. val leys are present not only in the landscape; they are also found buried in the subsurface. in denmark, buried valleys have oc ca sionally been described on the basis of borehole data and early geoelectrical methods (e.g. sorgenfrei & berthelsen 1954; lykke-andersen 1973; binzer & stock marr 1994). based on the large amount of newly collected hydrogeophysical data in denmark, it has recently become possible to define such valleys as tunnel valleys and to acknowledge their wide distribution in the subsurface (sandersen & jørgensen 2003; jørgensen et al. 2005; jør gensen & sandersen 2006). analysis of these data has re vealed dense networks of tunnel valleys and has significantly improved our understanding of their distribution, geometry and sedimentary infill. in the following, we review this work and identify new perspectives for the interpretation of the quaternary succession in denmark. the work was initiated in 1998 by the former danish counties (amter), and is currently continued by the geological survey of denmark and greenland and the ‘miljøcentre’ (environment centres). as part of this project, the buried valleys are continuously being mapped as new data are collected. data the spacing between and the quality of deep boreholes in denmark are generally insufficient to precisely delineate buried valleys (thomsen et al. 2004). borehole data are, however, important as they can often be used to verify the presence of valleys indicated from geophysical data, and provide valuable information about the sedimentary infill of the valleys. since the coverage of boreholes is generally too sparse for proper valley delineation, other types of data distributed in denser grids are needed. densely spaced data can be collected using the transient electro-magnetic (tem) geophysical method (sørensen & auken 2004) that gives deep penetramapping of buried tunnel valleys in denmark: new perspectives for the interpretation of the quaternary succession flemming jørgensen and peter b.e. sandersen © geus, 2008. geological survey of denmark and greenland bulletin 15, 33–36. available at: www.geus.dk/publications/bull 33 fig. 1. two selected areas with high densities of mapped buried valleys. a: north-western part of jylland. b: north-west of århus. aarhusårhus randers 8°30′e 9°40′e 10°10′e 56°50′n 56°20′n 9°00′e skive thisted struer lemvig mors buried valley town fig. 4b 10 km 10 km a a 56°30′n ba fig. 4c fig. 4a denmark b tion and an acceptable resolution. during the last ten years, this method has been extensively used in denmark, allowing many buried valleys to be mapped. high-resolution seismic data provide detailed 2d structural information of the valley architecture and allow detailed structural elements to be resolved, even in areas with poor lithological contrasts. however, data acquisition is costly and usually limited to widely spaced lines. the most effective method to map and investigate buried valleys is using a combination of the tem method, seismic data and borehole data (jørgensen et al. 2003). in the present project, all available and relevant lithological and geophysical data from the danish onshore area are being evaluated and examined in order to map and describe buried valleys (sandersen & jørgensen 2006). a set of criteria is used to obtain a high degree of certainty and objectivity in the valley delineation. the most important criterion is that the lateral extent and the orientation of the valleys must be unambiguously expressed in the available data. no interpolations between surveyed areas are made. valley characteristics more than 2500 km of buried valleys have currently been mapped in denmark using this procedure. the highest valley density is found in areas where tem data have been collected in dense grids and where the conditions for this specific method are ideal. such areas are found, for example, northwest of århus and in the north-western part of jylland (fig. 1). although many buried valleys have been mapped in these areas, even more are expected to exist because (1) not all valleys can be mapped with the methods in use, and (2) the study areas are not yet entirely covered by such data. only a minor part of the country has been surveyed by the tem method and all valleys are not represented in the collected data, so the maps can at best be considered to show the minimum occurrence of buried valleys. the general distribution and density of buried valleys therefore cannot be directly seen from the maps. thus, while a limited number of valleys have been mapped in western jylland, seismic data and borehole data indicate that many more buried valleys are present in this part of denmark than shown in fig. 2. delineation of buried valleys in western jylland is difficult due to low data coverage. the buried valleys can be divided into different generations that were formed during successive glaciations (jørgensen et al. 2005; jørgensen & sandersen 2006). the valleys often cross-cut each other and sometimes they have different preferred orientations. a complicated pattern with three to five generations of valleys can often be distinguished in areas with high data resolution (fig. 3). if the ice flows were parallel or nearly-parallel to pre-existing valleys, these valleys would be liable to repeated erosion and re-filling producing a complicated cut-and-fill setting within them. multiple generations of valley erosion are therefore also often found within the buried valleys themselves. the valley architecture is normally complex due to this cut-and-fill history, but also due to glaciotectonic disturbances. this complexity can be observed in both seismic and tem data (fig. 4). the complex fill is also often indicated by borehole data that show strong lithological variations over short distances. the valley fill is deposited in subglacial environments, in glacial lakes and rivers, in marine and terrestrial environments. statistical analyses of borehole data in all the mapped valleys show that by far the most common infill sediment types are tills and meltwater deposits (sandersen & jørgensen 2006). in this group, 52% are coarse meltwater deposits (sand and gravel), 30% are clay-rich till and 18% are finegrained meltwater deposits (silt and clay). the mapped buried valleys vary in depth with the deepest features exceeding 350 m from valley floor to shoulder (jør gensen & sandersen 2006). the width is generally between 0.5 and 1.5 km, but widths of up to 4 km occur. the lengths of the valleys are difficult to evaluate, because many of the areas surveyed so far are small. however, some valleys in larger survey areas exceed 25–30 km in length. some striking features of the valleys are that they often terminate abruptly and that they are highly irregular with depressions and thresholds along their floors. it is difficult to determine the ages of the buried valleys. the infill sediments can be dated in some cases and provide a minimum age for the time of valley erosion. however, based on the 34 denmark seismic line buried valley on seismic line mapped buried valley 8°30′e 5 km 56°00′n 8°45′e 9°00′e fig. 2. map of an area in western jyllland, showing the location of highresolution reflection seismic lines. the buried valleys cannot be accurately mapped from data obtained by this method alone, due to the wide line spacing. the valleys shown have mainly been mapped from data obtained by area-covering methods such as the tem method. modified from sandersen & jørgensen (2006). time gap between the age of the eroded sediments and the age of the infill sediments, buried valleys can be shown to have formed both during the weichselian, the saalian and the elsterian glacial stages (jørgensen & sandersen 2006). some of the valleys were probably formed during older glacial stages. valley formation the mapped buried valleys are comparable to open tunnel valleys found in the present-day danish landscape with respect to morphology and dimensions (jørgensen & sandersen 2006). most likely, they therefore formed as tunnel valleys and were subsequently buried by younger sediments. subglacial meltwater erosion is believed to have been responsible for most of the erosion, whereas direct glacial erosion may have contributed to the formation of the widest valleys. the presence of multiple generations of valleys shows that the area was repeatedly transgressed by glaciers during the pleistocene (jørgensen & sandersen 2006). tunnel valleys have a pronounced tendency to be re-used during repeated cycles of glacial erosion, which produces their cut-and-fill structure. the degree of re-use is presumably determined by differences in the morphology, erodibility and hydraulic conductivity of the substratum. however, the occurrence of proand subglacial permafrost may also have played an important role (jørgensen & sandersen 2006). consequences and perspectives buried valleys are widespread in the danish subsurface, occurring as complex, cross-cutting structures established during repeated erosion and infill. these characteristics require close attention when stratigraphic correlations are made, because discontinuous and complex geological settings are to be expected. 35 −100 −60 20 −20 60 −200 −80 −40 40 0 80 −120 elevation (m) 10 km 10°15′e10°00′e9°45′e10°15′e10°00′e9°45′e 56°15′n56°15′n a b fig. 3. a: map of an area north-west of århus showing the interpolated surface of a low-resistive bottom layer in tem soundings. large parts of the valleys have been eroded into this surface. b: cross-cutting valley generations found in the same area. black represents the oldest generation; light grey the youngest generation. modified from jørgensen & sandersen (2006). e le va tio n (m ) resistivity (ohm) 0 8040 200160120 0 40 80 −40 −80 −120 a b c 1000 m 2500 m 1000 m e le va tio n (m ) 0 40 −40 −80 −120 e le va tio n (m ) 0 40 −40 −80 −120 fig. 4. three profiles showing the complex geology related to buried valleys. a: a seismic profile across a buried valley with a cut-and-fill setting. interpretation shown by a red line. b: tem profile across a series of valleys. c: a tem profile across a valley with a cut-and-fill setting. interpretations on the tem profiles are shown with black lines. see fig. 1 for location of the profiles. modified from jørgensen et al. (2005) and jørgensen & sandersen (2006). correlation requires that the correlated sites are situated sufficiently close to each other so that they may not be significantly influenced by an unrecognised valley. the farther the valleys are apart, the more difficult correlation of their infill becomes. if one of the sites is situated inside an unrecognised or a poorly defined valley, stratigraphic correlations are likely to be erroneous, difficult or impossible to carry out. the discontinuous and complex geological setting of the valleys arises from the presence of unconformities produced by the valley erosion and from the limited lateral extent of the valleys and high variation of the sedimentary valley infill. unconformities complicate correlations of the incised layers, and the limited lateral valley extent and high variation of the infill complicate the correlation of the infill sediments. marked facies variations may make correlation of the infill sediments very difficult – even if the infill sediments were deposited simultaneously. the complications become even higher for valleys with cut-and-fill structures and when valleys cross-cut. erosion of the valleys implies that much of the sedimentary record has been removed, but as erosion and sedimentation often occurred during successive events, sedimentary remnants are occasionally left behind between the unconformities. therefore, the establishment of a coherent stratigraphy requires a dense data network and a large number of boreholes as well as a thorough knowledge of the outline and structure of the buried valleys and the relative age relationships between different valley generations in a given area. the challenges for stratigraphic correlations that arise from the presence of buried valleys as outlined above, also affect the construction of geological and hydrogeological models. since the valleys often cross-cut older layers and contain aquifers as well as aquitards, they have the ability to significantly influence the flow regime of groundwater (e.g. shaver & pusc 1992). furthermore, chances are high that buried valleys occur in almost all parts of denmark. in order to construct such models it is crucial to know where the valleys occur and understand how they were formed. while the occurrence of buried valleys imposes a series of new challenges, it may also provide new possibilities for the reconstruction of the pleistocene record. gradually increasing knowledge and overview of the erosional and depositional history of the valleys may lead to improved general understanding of the development of the pleistocene. it is likely that the deeper parts of some deep valleys may have withstood events of extensive glacial erosion during the pleistocene. thus some valleys may contain remnants of glacial or interglacial sediments, and in that case, buried valleys may contain sedimentary records from the early and early middle pleistocene. such records are currently very rarely reported from the onshore parts of denmark. acknowledgement the project is supported by the environmental centres in denmark. references binzer, k. & stockmarr, j. 1994: geological map of denmark, 1:500 000. pre-quaternary surface topography of denmark. danmarks geologiske undersøgelse kortserie 44, 10 pp., 2 maps. huuse, m. & lykke-andersen, h. 2000: over-deepened quaternary valleys in the eastern danish north sea: morphology and origin. quat ernary science reviews 19, 1233–1253. jørgensen, f. & sandersen, p.b.e. 2006: buried and open tunnel valleys in denmark – erosion beneath multiple ice sheets. quaternary science reviews 25, 1339–1363. jørgensen, f., lykke-andersen, h., sandersen, p.b.e., auken, e. & normark, e. 2003: geophysical investigations of buried quaternary valleys in denmark: an integrated application of transient electromagnetic soundings, reflection seismic surveys and exploratory drillings. journal of applied geophysics 53, 215–228. jørgensen, f., sandersen, p.b.e., auken, e., lykke-andersen, h. & sørensen, k. 2005: contributions to the geological mapping of mors, denmark – a study based on a large-scale tem survey. bulletin of the geological society of denmark 52, 53–75. lykke-andersen, h. 1973: en begravet dal i præ-kvartæret ved århus. dansk geologisk forening årsskrift for 1972, 111–118. madsen, v. 1921: terrainformerne på skovbjerg bakkeø. danmarks geo logiske undersøgelse iv. række 1(12), 24 pp. o’cofaigh, c. 1996: tunnel valley genesis. progress in physical geography 20, 1–19. sandersen, p.b.e. & jørgensen, f. 2003: buried quaternary valleys in western denmark – occurrence and inferred implications for groundwater resources and vulnerability. journal of applied geophysics 53, 229–248. sandersen, p.b.e. & jørgensen, f. 2006: kortlægning af begravede dale i jylland og på fyn. opdatering 2005–2006, 201 pp. de jysk-fynske amters grundvandssamarbejde. vejle: vejle amt & watertech a/s. shaver, r.b. & pusc, s.w. 1992: hydraulic barriers in pleistocene buriedvalley aquifers. ground water 30, 21–28. smed, p. 1998: die entstehung der dänischen und norddeutschen rin nentäler (tunneltäler) – glaziologische gesichtspunkte. eiszeitalter und gegenwart 48, 1–18. sørensen, k.i. & auken, e. 2004: skytem – a new high-resolution helicopter transient electromagnetic system. exploration geophysics 35, 191–199. sorgenfrei, t. & berthelsen, o. 1954: geologi og vandboring. danmarks geologiske undersøgelse iii. række 31, 106 pp. thomsen, r., søndergaard, v.h. & sørensen, k.i. 2004: hydrogeological mapping as a basis for establishing site-specific groundwater protection zones in denmark. hydrogeology journal 12, 550–562. ussing, n.v. 1907: om floddale og randmoræner i jylland. oversigt over det kongelige danske videnskabernes selskabs forhandlinger 4, 161–213. 36 authors’ addresses f.j., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: flj@geus.dk p.b.e.s., grontmij | carl bro a/s, dusager 12, dk-8200 århus n, denmark.. geological survey of denmark and greenland bulletin 20, 2010, 47–50 47 diagenetic precipitation of silicate minerals such as quartz and clay minerals can reduce the permeability and porosity of chalk as they precipitate as cement in pores (e.g. taylor & lapré 1987; maliva & dickson 1992). however, the precipitation can also result in early lithification and help to preserve porosity during burial. several studies of diagenesis have been carried out on chalk samples from the north sea oil fields (e.g. scholle 1977; maliva & dickson 1992; hancock 1993) but only a few have focused on silica diagenesis (fabricius & borre 2007; fabricius et al. 2007). silica is not a major constituent of chalk but is abundant in some intervals where it influences the reservoir properties, as seen for example in the ekofisk formation and in a few intervals in the underlying tor formation in the eldfisk and ekofisk fields. this distribution may indicate that the variations are linked to facies and palaeo-oceanography; the silica either representing diagenetically reprecipitated biogenic silica, volcanic ash falls or a high input of detrital minerals (scholle 1977; kennedy 1987; fabricius & borre 2007). this paper discusses diagenesis of quartzand kaoliniterich intervals in the deeply buried upper maastrichtian chalk of the upper tor formation (ta layer) in core 2/7-b-12 a from the eldfisk field in the norwegian sector of the north sea (fig. 1). the study was carried out as part of a phd project at the geological survey of denmark and greenland and the department of geography and geology, university of copenhagen (madsen 2009). the objective of the study was to understand the processes leading to formation and enrichment of quartz and kaolinite and the resulting influence on porosity. lithology and mineralogy the upper tor formation consists of grey bioturbated chalk with stylolites, argillaceous solution seams and a few marly layers in the studied core (fig. 2). a single firmground is present at 10 480 ft. the chalk is usually almost pure carbonate with 96–97% calcite. the insoluble residue mainly consists of quartz and kaolinite, with small amounts of illite-smectite, feldspar, fluor-apatite, crandelite (caal3(po4)1.5(oh)·5h2o), dolomite, pyrite and fluorite. however, up to 12% insoluble residue is present in two intervals at 10 420–10 431 and 10 470–10 480 ft. the 10 420– 10 431 ft interval contains abundant stylolites, and overall the insoluble residue consists of the same minerals as in the pure chalk. the 10 470–10 480 ft interval only contains few marly layers and stylolites, and the insoluble residue in this interval is dominated by kaolinite and quartz. quartz occurs as euhedral crystals, aggregates and subhedral crystals. the euhedral crystals are six-sided and doubleterminated by six-faced pyramids up to 10 μm long (fig. 3a). both crystals and aggregates are commonly associated with kaolinite. the quartz crystals commonly sit on kaolinite crystals but may also enclose the kaolinite (fig. 3c). in the 10 420–10 431 ft interval the subhedral quartz crystals, up to 25 μm in size, enclose recrystallised coccolith fragments and older quartz crystals. the subhedral quartz crystals only show few well-developed crystal faces, commonly with imprints from recrystallised coccolith fragments. kaolinite occurs as hexagonal crystals arranged in booklets, commonly around 5 μm in diameter (fig. 3a–c). the kaolinite is found both in voids and as part of the matrix. it is often associated with quartz aggregates and single quartz crystals (fig. 3a, b). energy-dispersive x-ray spectroscopy silica diagenesis and its effect on porosity of upper maastrichtian chalk – an example from the eldfisk field, the north sea heine buus madsen © geus, 2010. geological survey of denmark and greenland bulletin 20, 47–50. open access: www.geus.dk/publications/bull fig. 1. map of the north sea region showing structural elements and the location of the eldfisk field. dashed lines: national borders. 0° 8° 58°n 50°n germany united kingdom north sea denmark eldfisk field central graben 250 km norway sweden 4848 analyses show that the kaolinite primarily consists of si, al and o. illite-smectite is present in trace amounts in all samples analysed by x-ray diffraction. it has a delicate, flaky morphology (fig. 3d) and contains small amounts of k and mg in addition to si, al and o. petrophysical characteristics of intervals rich in insoluble residue in order to determine the petrophysical properties of the intervals high in insoluble residue, petrophysical logs, lithological variations and mineralogy were compared with each other. the interval 10 420–10 431 ft contains much more quartz and kaolinite than the surrounding chalk. the lower part of the interval shows high gamma-ray values (gr; fig. 2). the neutron log (nphi) and density log (rhob) responses are similar to the surrounding chalk. high gr values are also seen below the 10 420–10 431 ft interval, indicating that the interval rich in insoluble residues extends down to 10 440 ft. in the 10 470–10 480 ft interval high gr values correlate with high contents of kaolinite and quartz, whereas no obvious peaks in nphi and rhob can be related to changes in the mineralogy (fig. 2). a neutron log versus density log plot of the studied succession shows that the chalk mainly plots between the sandstone and limestone trends (fig. 4). this is due to the presence of hydrocarbons which are lighter than saline pore water. both intervals rich in insoluble residue and the chalk between 10 415–10 485 ft have higher density and lower porosity than the upper interval at 10 400–10 415 ft. the estimated porosity of the chalk in the 10 400–10 415 ft interval is around 35% whereas the quartzand kaolinite-rich intervals have porosities around 25% (fig. 4). the high gammaray values (dark-brown plus signs) represent clay-rich parts and have estimated porosities of 20–25%. silicate diagenesis the quartz and kaolinite may either have derived from detrital grains or from devitrification of volcanic ash. however, the latter produces bentonite which is dominated by illite and montmorillonite. hence, volcanic ash cannot be the source. weathering products such as smectite, kaolinite, quartz and feldspar from the nearby baltic shield are a more likely source for the quartz and kaolinite as also suggested by kennedy (1987). the distribution of quartz and kaolinite in the two insoluble-rich intervals also indicates an event of longer duration than a volcanic eruption. however, some of the authigenic quartz probably originated from phase transformations of biogenic opal-a to opal-ct and finally to quartz. the occurrence of long booklets of kaolinite indicates that some of it precipitated from solution (fig. 3b). but the close association between quartz and kaolinite seen in the scanning electron microscope images indicates that they originated from mineral reactions (fig. 3a, c). the disappearance of smectite and the formation of illite with depth are well known in the central graben in the north sea and in other sedimentary basins (compton 1991; abercrombie et al. 1994; drits et al. 1997). however, since kaolinite is the most abundant clay mineral in the studied interval, a slightfig. 2. lithological and petrophysical logs of the upper part of the tor formation, showing the 10 400–10 495 ft interval in the 2/7-b-12 a core. chalk quartz kaolinite apatite dolomite pyrite feldspar crandelite argillaceous chalk stylolite marl layer insoluble residue (%) 0.6 0 1.7 2.7 0 gr (api) rhob (g/cm3) nphi (v/v decimal) 50 10 440 10 445 10 450 10 455 10 460 10 465 10 470 10 475 10 480 10 485 10 425 10 420 10 430 10 435 10 415 10 410 10 405 10 400 depth (ft) 0 2 4 6 8 10 12 49 ly different mineral reaction is suggested where smectite is mainly transformed to kaolinite, following the equation: smectite + k-feldspar → kaolinite + quartz + illite-smectite the formation of kaolinite and illite-smectite is probably the result of release of mg and fe from the octahedral layer and si from the tetrahedral layer in the smectite (hower et al. 1976), which enables precipitation of quartz near the kaolinite as seen in fig. 3. during this transformation aluminium is preserved in the structure (compton 1991). the released mg and fe must have diffused away because only small amounts of dolomite or other mgand fe-rich minerals are observed in the chalk (fig. 2). alternatively the smectite was originally poor in mg and fe. additional potassium, from k-feldspar or other sources is needed to form smectite-illite (hoffman & hower 1979; compton 1991). the fact that kaolinite is the most common authigenic clay mineral indicates that detrital potassiumrich minerals were rare. however, dissolution of potassiumrich minerals has probably contributed with al, si and some k to the formation of kaolinite and smectite-illite, with excess si precipitating as quartz. the mineral assemblage with kaolinite and quartz indicates a potassium-poor system, but the high gamma-ray values recorded in the two insoluble-rich intervals imply that minerals containing potassium or other radiogenic elements are indeed present (fig. 2). porosity reduction due to a high primary content of detrital smectite and illitesmectite in the 10 420–10 431 and 10 470–10 480 ft intervals, the primary sorting and hence the initial porosities were lower than in the surrounding chalk. the neutron log versus density log plot (fig. 4) shows an estimated porosity of c. 35% for samples of pure chalk (the 10  400–10  415 ft interval), and porosities of c. 27% and 22% for the 10 420–10 431 and 10 470–10 480 ft intervals that are rich in insoluble residue. this confirms that the content of detrital minerals had a great effect on the initial porosities. however, precipitation of quartz and kaolinite in pores reduced the porosity further. conclusions two intervals with high contents of insoluble residue, up to 12%, were found in the upper tor formation in the 2/7-b-12 a core from the eldfisk field. the insoluble residue primarily consists of authigenic quartz and kaolinite with subordinate illite-smectite. the kaolinite, illite-smectite and quartz formed both by precipitation from solution and by mineral reactions during burial. the release of mg, fe and si and the addition of al in the crystal latice of detrital smectite resulted in a mineral reaction that formed most of the kaolinite. dissolution of feldspar contributed al, and the released si precipitated as quartz at or near the kaolinite. the intervals represent potassium-poor systems where kaolinite precipitatq k k q q is is k q q q k a b 1 μm 5 μm5 μm k a b dc 2 μm 5 μm 1 μm2 μm fig. 3. scanning electron microphotographs of chalk samples illustrating components of the insoluble residue. a: double-terminated quarts crystal (q) sitting on kaolinite (k). b: quartz aggregates (lower left corner (q), consisting of sub-micron sized quartz crystals) and booklets of kaolinite (k). c: kaolinite (k) embedded in quartz (q). d: flaky illite-smectite (is). 5050 ed rather than illite-smectite. the high content of insoluble residue in the 10 420–10 431 ft and 10 470–10 480 ft intervals was probably caused by a high content of detrital minerals, causing poor sorting and lower porosities compared to the more pure chalk. precipitation of kaolinite and quartz caused a minor additional reduction of the porosity in the insoluble residue-rich intervals. acknowledgements conocophillips norway is thanked for financial support and for providing core material and geophysical logs. references abercrombie, h.j., hutcheon, i.e.. bloch, j.d. & decaritat, p. 1994: silica activity and the smectite-illite reaction. geology 22, 539–542. compton, j.s. 1991: origin and diagenesis of clay minerals in the monterey formation, santa maria basin area, california. clays and clay minerals 39, 449–466. drits, v.a., sakharov b.a., lindgreen, h. & salyn, a. 1997: sequential structure transformation of illite-smectite-vermiculite during diagenesis of upper jurassic shales from the north sea and denmark. clay minerals 32, 351–371. fabricius, i.l. & borre, m.k. 2007: stylolites, porosity, depositional texture, and silicates in chalk facies sediments. ontong java plateau – gorm and tyra fields, north sea. sedimentology 54, 183–205. fabricius, i.l., røgen, b. & gommesen, l. 2007: how depositional texture and diagenesis control petrophysical and elastic properties of samples from five north sea chalk fields. petroleum geoscience 13, 81–95. hancock, j.m. 1993: the formation and diagenesis of chalk. in: downing, r.a., price, m. & jones, g.p. (eds): the hydrogeology of the chalk of north-west europe, 14–34. new york: oxford university press. hoffman, j. & hower, j. 1979: clay mineral assemblages as low grade metamorphic geothermometers: application to the thrust faulted disturbed belt of montana, u.s.a. in: scholle, p.a. & schluger, p.r. (eds): aspects of diagenesis. society of economic paleontologists and mineralogists (sepm) special publications 26, 55–79. hower, j., eslinger, e.v., hower, m. & perry, e.a. 1976: the mechanism of burial metamorphism of argillaceous sediments: 1. mineralogical and chemical evidence. geological society of america bulletin 87, 725–737. kennedy, w.j. 1987: sedimentology of late cretaceous–paleocene chalk reservoirs, north sea central graben. in: brooks, j. & glennie, k.w. (eds): petroleum geology of north west europe, 469–481. proceedings of the 3rd conference on petroleum geology of north west europe. london: graham & trotman. madsen, h.b. 2009: microbial and burial induced silica diagenesis in upper cretaceous – paleogene chalk in denmark and the north sea. phd thesis, university of copenhagen, denmark. danmarks og grønlands geologiske undersøgelse rapport 2009/41, 33 pp. maliva, r.g. & dickson, j.a.d. 1992: microfacies and diagenetic controls of porosity in cretaceous/tertiary chalks, eldfisk, norwegian north sea. aapg bulletin 76, 1825–1838. tulsa, oklahoma: american association of petroleum geologists. scholle, p.a. 1977: chalk diagenesis and its relation to petroleum exploration. aapg bulletin 61, 982–1009. tulsa, oklahoma: american association of petroleum geologists. taylor, s.r. & lapré, j.f. 1987: north sea chalk diagenesis: its effect on reservoir location and properties. in: brooks, j. & glennie k.w. (eds): petroleum geology of north west europe, 483–495. proceedings of the 3rd conference on petroleum geology of north west europe. london: graham & trotman. 0 5 10 15 20 25 30 35 40 45 0 5 10 15 20 30 35 40 45 0 5 10 15 20 25 30 35 r h o b ( g/ cm 3 ) nphi (v/v decimal) sandstone limestone dolomite kaolinite illite 1.8 2.0 2.2 2.4 2.6 2.8 -0.05 0.05 0.15 0.25 0.35 0.45 z axis: gamma ray (api) 5 11 17 23 29 35 10 400–10 415 ft 10 470–10 480 ft 10 420–10 431 ft fig. 4. nphi versus rhob plot of the 10 400–10 495 ft interval in core 2/7-b-12 a from the eldfisk field. trends for sandstone, limestone and dolomite porosities are shown for comparison. the pure chalk at 10 400– 10  420 ft mostly shows high estimated porosities (c. 35%), whereas the lower section and the intervals rich in insoluble residue show lower porosities. the dark brown coloured data points reflect high gammaray value samples, which also have the lowest estimated porosities (the 10 470–10 480 ft interval). author’s address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: heinebuus@yahoo.com geological survey of denmark and greenland bulletin 17, 2009, 1-8 geological survey of denmark and greenland ministry of climate and energy geological survey of denmark and greenland bulletin 17 • 2009 review of survey activities 2008 edited by ole bennike, adam a. garde and w. stuart watt rosa_2008:rosa-2008 01/07/09 15:47 side 1 geological survey of denmark and greenland bulletin 17 keywords geological survey of denmark and greenland, survey organisations, current research, denmark, greenland. cover photographs from left to right 1. inspection of cores in jameson land, east greenland. photo: john boserup. 2. field work in west greenland. photo: denis schlatter. 3. investigations and sampling of an outcrop at conde, bahia, brazil. photo: peter japsen. 4. field experiments with remediation of contaminated soil in vadsby, west of copenhagen. photo: knud erik s. klint. frontispiece: facing page visit at the automatic weather station ‘lower nuuk’ on 30 july 2008 for data retrieval and maintenance. the station is located on the gla cier qamanaarsuup sermia near the margin of the greenland ice sheet, east of godthåbsfjord (64º29´n, 49°31´w). it collects weather information and data about the local ablation and was established in 2007 as part of the promice project (www.promice.org). photo: søren nielsen. chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, department of earth sciences, university of aarhus; minik rosing, geological museum, university of copenhagen; finn surlyk, department of geography and geology, university of copenhagen scientific editors: ole bennike, adam a. garde and w. stuart watt editorial secretaries: jane holst and esben w. glendal referees: (dk = denmark; numbers refer to first page of reviewed article): anonymous (29, 29, 33, 45, 49, 77, 77, 81); lars christiansen, dk (41); finn dalhoff, dk (13); gregers dam, dk (53); edward f. duke, usa (69); ida fabricius, dk (17); henrik friis, dk (25); mariejosé gaillard, s (37); ole graversen, dk (81); karen hanghøj, usa (61); bent hasholt, dk (45); jens havskov, n (9); ole humlum, n (74); paul martin holm, dk (61); jens konnerup-madsen, dk (65); john a. korstgård, dk (57); nicolaj krog larsen, s (33); poul-henrik larsen, dk (53); christoph mayer, g (74); florence mazier, s (37); klaus mosegaard, dk (9); john myers, australia (49); allan aasbjerg nielsen, dk (69); henrik olsen, dk (41); graham pearson, uk (65); asger ken pedersen, dk (57); gunver krarup pedersen, dk (21); jan audun rasmussen, dk (21); martin sønderholm, dk (25); morten gjetting stage, dk (13); ole v. vejbæk, dk (17). illustrations: stefan sølberg, with contributions from benny m. schark lay-out and graphic production: henrik klinge pedersen and annabeth andersen printers: schultz grafisk, albertslund, denmark manuscripts submitted: 23 january – 27 march 2009 final versions approved: 25 may 2009 printed: 8 july 2009 issn 1603-9769 (review of survey activities) issn 1604-8156 (geological survey of denmark and greenland bulletin) isbn 978-87-7871-250-9 citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 17, 84 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk © de nationale geologiske undersøgelser for danmark og grønland (geus), 2009 for the full text of the geus copyright clause, please refer to www.geus.dk/publications/bull rosa_2008:rosa-2008 01/07/09 15:47 side 2 rosa_2008:rosa-2008 01/07/09 15:47 side 3 4 7 review of survey activities 2008 f.g. christiansen 9 earthquake in southern sweden wakes up denmark on 16 december 2008 p.h. voss, t.b. larsen, l. ottemöller and s. gregersen 13 the potential for large-scale, subsurface geological co2 storage in denmark p. frykman, l.h. nielsen, t. vangkilde-pedersen and k.l. anthonsen 17 increased oil recovery from halfdan chalk by flooding with co2-enriched water: a laboratory experiment d. olsen 21 ladinian palynofloras in the norwegian– danish basin: a regional marker reflecting a climate change s. lindström, h. vosgerau, s. piasecki, l.h. nielsen, k. dybkjær and m. erlström 25 fingerprinting sediments along the west coast of jylland: interpreting provenance data c. knudsen, t. kokfelt, t. aagaard, j. bartholdy and m. pejrup 29 structural development of maglevandsfald: a key to understanding the glaciotectonic architecture of møns klint, se denmark s.a.s. pedersen and p. gravesen 33 fracture valleys in central jylland – a neotectonic feature p.r. jakobsen and s.a.s. pedersen 37 soil erosion and land-use change during the last six millennia recorded in lake sediments of gudme sø, fyn, denmark p. rasmussen and j. olsen thailand laos pdr ghana uganda tanzania mozambique vietna greenland kenya spain chile germany the netherlands denmark brazil zambia yemen cambodia mali sudan bahrain geus working areas 2008. orange areas are covered in this volume. for further information on other working areas please refer to our website: www.geus.dk/international rosa_2008:rosa-2008 01/07/09 15:47 side 4 5 41 geophysical methods and data administra tion in danish groundwater mapping i. møller, v.h. søndergaard and f. jørgensen 45 water budget of skærsø, a lake in south-east jylland, denmark: exchange between groundwater and lake water b. nilsson, p. engesgaard, j. kidmose, s. karan, m.c. looms and m.c.s. frandsen 49 geological observations in the southern west greenland basement from ameralik to frederikshåb isblink in 2008 n. keulen, a. scherstén, j.c. schumacher, t. næraa and b.f. windley 53 shallow core drilling and petroleum geology related field work in east and northeast greenland 2008 j.a. bojesen-koefoed, m. bjerager and s. piasecki 57 the bedrock geology under the inland ice: the next major challenge for greenland mapping p.r. dawes 61 developing a 3-d model for the skaergaard intrusion in east greenland: constraints on structure, mineralisation and petrogenetic models t.f.d. nielsen, s.d. olsen and b.m. stensgaard 65 diamonds and lithospheric mantle properties in the neo proterozoic igneous province of southern west greenland a. steenfelt, s.m. jensen, t.f.d. nielsen, k.k. sand and k. secher 69 using spectral mixture analysis of hyperspectral remote sensing data to map lithology of the sarfartoq carbona tite complex, southern west greenland e. bedini and t. tukiainen 73 glaciological investigations at the malmbjerg mining prospect, central east greenland m. citterio, r. mottram, s.h. larsen and a. ahlstrøm 77 holocene climate variability in southern greenland: results from the galathea 3 expedition n. nørgaard-pedersen, n. mikkelsen, m.d. poulsen and a.s. simonsen 81 post-rift landscape development of northeast brazil j.m. bonow, p. japsen, p.f. green, p.r. cobbold, a.j. pedreira, r. lilletveit and d. chiossi thailand laos pdr uganda tanzania mozambique vietnam kenya phillipines mark zambia yemen cambodia sudan bahrain rosa_2008:rosa-2008 01/07/09 15:47 side 5 rosa_2008:rosa-2008 01/07/09 15:47 side 6 7 following a number of years with major changes of the scientific environment in denmark and also within the management of the geological survey of denmark and greenland (geus), 2008 was a year of stability and consolidation, a situation that will hopefully continue. many new projects have been initiated and many previous projects have been completed at a time with strong focus on geus’ activities politically, commercially and from the media. this sixth annual issue of review of survey activities describes selected projects that geus and its partners carry out in denmark, greenland and internationally. together with the previous five published issues (also available at www.geus.dk), it provides a good overview of the survey’s many different types of research and advisory activities. it contains a total of 19 four-page papers: ten on denmark, eight on greenland and one on international work. geology was on the lips of most of the danish population an early morning in december 2008 when one of the strongest earthquakes recorded in scandinavia woke up hundreds of thousands of people in southern sweden and on sjælland. one paper in this issue describes the background and details of the earthquake, including input from the public that contacted geus through its website. reduction of the emission of co2 is high on the political agenda in denmark and internationally. one of the possibilities to reduce co2 emission from large point sources is to use carbon capture and storage (ccs). the danish subsurface has a high potential to store co2, and structures such as the vedsted structure in northern jylland could be among the first dozen storage facilities utilised in europe, and thereby become a key area for detailed research and monitoring for many years to come. the background for ccs and geological possibilities in denmark are described in one paper. oil and gas exploration and production are still very important for the economy of denmark, and geus has a strong emphasis on research within this field. two papers concentrate on petroleum geology. one of them is based on laboratory flooding experiments and describes the possibility of increasing oil recovery from reservoirs in chalk using injection of co2-enriched water; the other provides a detailed biostratigraphic correlation of the late triassic succession in the norwegian–danish basin. most surface features in denmark have been formed by glacial and coastal processes during the quaternary. several papers in this issue describe such processes; one of them demonstrates the use of sophisticated analytical techniques such as computer-controlled, scanning electron microscopy of heavy minerals and laser ablation, inductively coupled mass spectro metry of zircon grains to describe erosion and re-deposition of sand along the west coast of jylland. the structural development of the famous møns klint geosite is dealt with in one paper, and the occurrence of neotectonic fracture valleys in central jylland in another. a third paper describes soil erosion and land use change during the last six millennia as recorded in lake sediments from gudme sø on fyn. groundwater mapping and management have a very high priority in denmark. one paper describes the many different geophysical methods that are used in hydrogeological mapping, as well as the administration of the geophysical data that are archived in a major database hosted at geus. another paper describes the exchange between lake water and groundwater of lake skærsø in jylland. in 2008 there was a high level of field activities in green land. the two largest campaigns in southern west green land and eastern greenland are described in individual papers. the west greenland field work was a follow-up on earlier projects focused on updating previous maps and thereby creating a better understanding of the potential distribution of mineral occurrences. the field work in eastern greenland is the start of a major oil industry sponsored programme that has been launched to support and promote petroleum exploration within the coming five-year period. it includes shallow core drilling. one paper addresses a question often raised by scientists and explorers: what is the bedrock geology under the inland ice that covers 81% of the total area of greenland? available review of survey activities 2008 flemming g. christiansen deputy director © geus, 2009. geological survey of denmark and greenland bulletin 17, 7–8. available at: www.geus.dk/publications/bull rosa_2008:rosa-2008 01/07/09 15:47 side 7 geological and geophysical data are shortly reviewed and ideas for future studies presented. another paper introduces a 3-d modelling of one of the best studied intrusions in the world, the paleocene skaergaard intrusion in east greenland. identification of significant platinum group and gold occurrences in this intrusion has lead to detailed investigations and exploration drilling over many years. the last decade of diamond exploration in west greenland has provided a wealth of data on the dykes of kimberlite and ultramafic lamprophyres that may host diamonds. one paper summarises new data on petrology and age distribution of the dykes that have important implications for future diamond exploration. in a large country like greenland, use of remote sensing data is important and cost-effective in mapping and exploration. one paper presents an analysis of hyperspectral data from the sarfartoq carbonatite complex in west greenland. the paper illustrates that such data can be applied to mapping of individual rock types. the most recent processes and climate development in greenland are described in two papers of which one stresses the importance of applied glaciology for the exploitation of the malmbjerg molybdenum deposit in east greenland, because the site of a possible future mine is located between two glaciers. future access to the mining site and removal of ore and waste rock are highly dependent on the movement of the glaciers. the second paper addresses holocene climate variation and marine history in south greenland, based on a number of samples collected in bredefjord and narsaq sund during the galathea 3 expedition. a final paper on landscape development in brazil employs the same methods of landscape mapping and apatite fission track analysis that were used to study the uplift history of the margins of greenland and scandinavia. 8 rosa_2008:rosa-2008 01/07/09 15:47 side 8 geological survey of denmark and greenland bulletin 13, 2007, 65-68 65 experience and results from the danish groundwater monitoring programme that has been carried out systematically since 1990, have been used in a co-operative project between latvia and denmark. the main objective of the project was to obtain more detailed knowledge of the shallow latvian groundwater, to optimise the latvian groundwater monitoring programme and to support the implementation of euro pean legislation such as the water framework directive, the nitrate directive and the groundwater directive in latvia. comprehensive summaries describing the methodology of groundwater quality monitoring as well as the major results from the danish groundwater monitoring network can be found in geus (2005) and stockmarr (2005). until recently only few data on latvian groundwater quality were available, but in a project running from 2003 to 2006, 800 samples from groundwater, springs and drains have been analysed for a large number of components resulting in a comprehensive overview of the status of latvian groundwater (fig. 1; gosk et al. 2006). the project agricultural influence on groundwater in latvia was carried out by the state geological survey of latvia and the geological survey of denmark and greenland (geus) and was supported by the danish environmental protection agency within the framework of the dancee programme (danish co-operation for environment in eastern europe). as a spin-off of the project this paper compares groundwater quality in the two countries. groundwater use and agricultural load drinking-water production in denmark is almost entirely based on groundwater abstracted from both shallow and moderately deep quaternary and pre-quaternary aquifers covered by more or less protective layers of till of variable thickness. the annual groundwater recharge is relatively large in denmark and is on an average, national scale sufficient to cover needs (henriksen & sonnenborg 2003). however, © geus, 2007. geological survey of denmark and greenland bulletin 13, 65–68. available at: www.geus.dk/publications/bull shallow groundwater quality in latvia and denmark edmund gosk, igors levins and lisbeth flindt jørgensen fig. 1. sampling sites of latvian groundwater. abstraction exceeds recharge in some areas, especially around the largest cities. furthermore, diffuse groundwater contamination by pesticides and nitrate has become an increasing problem since the middle of the last century, and a number of water-supply wells have been closed. up to now it has been possible to find new, clean groundwater resources, either by drilling deeper wells or by moving wells to less impacted areas. more than 60% of the danish territory is defined as arable land. the impact of agriculture on water quality has been known for many years, and measures aiming at a reduction of nutrient load started in the 1980s with the adoption of the first action plan on the aquatic environment. at that time the average nitrogen load had reached 130 kg n/ha, which resulted in a high risk of groundwater contamination and in actual quality problems in some areas of the country. drinking-water production in latvia is also primarily based on groundwater. only riga water supply uses a substantial amount of surface water to produce about half of the drinking water consumed in the latvian capital. the thick till layer that covers most of latvia is characterised by poor aquifer conditions where only little groundwater abstraction is possible from small, local water-bearing lenses. the centralised drinking-water supply in latvia is therefore typically based on confined aquifers screened at depths greater than 50 m. in areas with intensive agriculture (the southern and central parts of the country) the average depth of wells exceeds 100 m. due to the dramatic political and economical changes in latvia at the beginning of the 1990s groundwater abstraction decreased by a factor of 2–3 resulting in a significant, still progressing recovery of groundwater levels in the central and south-western parts of the country. in the remaining parts of the country groundwater levels have never been affected by abstraction. during the same period, the acreage of arable crops, the use of mineral fertilisers, the number of livestock and other parameters related to agricultural load, decreased by a factor of 3–10 (stålnacke et al. 2003). at present only 13% of the latvian area is under cultivation, and the average crop yield is about half of the danish figures. groundwater monitoring and knowledge of diffuse contamination due to the pressure from diffuse sources, groundwater monitoring plays an important role in water management and protection of groundwater resources in denmark. more than 6000 drinking-water supply wells, a network of more than 1400 specially selected screens in 70 groundwater monitoring areas, as well as five agricultural watershed monitoring areas equipped with a total of about 100 shallow screens (stockmarr 2005), together form the basis of the danish groundwater monitoring system. the detailed groundwater quality monitoring includes analyses of 97 elements, comprising 14 heavy metals and 34 pesticides and their metabolites (neri 2005). the regional groundwater quality monitoring network in latvia, established in the 1970s and 1980s, is based on 150 wells located within 45 monitoring stations. this network was designed to control depression cones (quantitative monitoring) rather than to monitor the various forms of diffuse pollution (qualitative monitoring). monitoring wells are typically screened in deep aquifers used for abstraction. until recently, the parameters analysed comprised major ions and nitrogen compounds, but no analyses of heavy metals and pesticides were made. groundwater quality data from abstraction wells consisted mainly of information collected during drilling. monitoring of drinking-water quality in latvia is almost entirely limited to sampling of tap water. although some of the larger water-supply companies collect samples from abstraction wells, the centralised data processing system for this type of analysis is not yet established. the amount of groundwater data recently collected in latvia is much smaller than comparable data collected in denmark. however, a large amount of information on the youngest groundwater in shallow aquifers has been collected in the period 2003–2005 as part of the present project. about 800 groundwater samples were analysed for a wide spectrum of elements, including heavy metals and pesticides (gosk et al. 2006). the determination of background concentration of various elements was based on data from shallow wells located outside local point-pollution plumes and urbanised and industrial areas. nitrates and pesticides as expected, the distribution of nitrates in danish and latvian groundwater differs considerably (fig. 2a). in the shallowest groundwater, the concentration of nitrates exceeds the quality standards of drinking water (50 mg/l no3) in 15% and 3.5% of the monitoring screens in danish and latvian wells, respectively. furthermore, the decrease of nitrate concentration with depth is much faster in the latvian than in the danish groundwater. the differences observed are caused by a lower agricultural load in latvia as well as by a higher denitrification potential and confinement degree of latvian aquifers and aquitards (fig. 2b). a statistical comparison of the groundwater residence times is impossible due to the very limited number of 66 age determinations of latvian groundwater (27 water samples were analysed by the cfc method). however, it seems that latvian groundwater is about twice as old as danish groundwater from the same depth. the recent investigations in latvia revealed some unexpected results with respect to nitrate in shallow aquifers. latvian springs capturing the youngest and the most mobile part of groundwater flow seem to be characterised by higher average nitrate concentrations compared to shallow wells drilled to the same aquifers (fig. 3). even springs located within areas without obvious nitrate load show persistently higher nitrate concentrations than neighbouring wells. after recognition of this phenomenon several high-yielding springs have been included in the groundwater monitoring network to better understand flow patterns and problems of diffuse nitrate pollution. investigations of contamination of latvian groundwater by pesticides have only just started, and the extent of this problem is not fully clear. vulnerable areas were identified in the project, and 111 groundwater samples from wells in these areas were analysed for a number of expected pesticides. in 30% of the samples, concentrations of pesticides exceeded the quality standards of drinking water (0.1 µg/l). this figure is identical to the frequency in shallow danish monitoring screens (stockmarr 2005). in denmark the most often found pesticide is bam (2,6-dichlorobenzamide), a metabolite from dichlobenil and chlorothiamide that prior to 1997 was commonly used as a total herbicide in urban areas, along roads, railways and at farm yards. the triazine group is the second most found pesticide group, but in latvia this group comes in third place, after the chlorophenoxy acids and trichloro acetic acid (tca); the latter was the most widely utilised herbicide during the soviet time and a by-product from the chemical industry. as the investigation was focused on potentially vulnerable areas and in the uppermost aquifers, an estimate of 30% occurrence of pesticides is highly pessimistic if the entire latvian groundwater is considered. taking into account the smaller acreage of treated areas and a higher confinement degree of main aquifers, the risk of contamination of water-supply wells by pesticides in latvia should be significantly lower than in denmark. however, as the pesticides were mainly used in the 1980s they may not yet have reached the main latvian aquifers containing groundwater older than 50 years. minor elements in order to compare danish and latvian shallow groundwater chemistry, distribution curves of twelve selected minor elements are presented in fig. 4. shallow latvian groundwater is characterised by higher background concentrations of aluminium and fluoride and lower concentrations of bromide, lithium, manganese, phosphorus and strontium when compared to the danish data. a majority of the observed differences reflect the difference in basic groundwater chemistry: the median values of the chloride, sulphate and permanganate indices in the danish data set (30, 45 and 3.1 mg/l, respectively) are significantly higher than the corresponding values for the latvian data set (10, 20 and 1.4 mg/l, respectively). 67 fig. 2. a: distribution of nitrate in danish and latvian wells compared to screen depth. danish data from 719 monitoring screens averaged for the period 1998–2003. latvian data from 717 monitoring and water-supply wells averaged for the period 2000–2005. b: distribution of dissolved oxygen in danish and latvian wells as a function of screen depth. danish data from 706 monitoring screens are averaged for the period 1998–2003. latvian data from 160 monitoring and investigative screens are averaged for the period 2003–2005. fig. 3. distribution of nitrate in 85 springs and 358 shallow wells within latvian agricultural areas. 68 the anoxic and neutral conditions in the majority of latvian aquifers are unfavourable for migration of most heavy metals, whose concentrations are well below the quality standards of drinking water. exceptions are manganese and arsenic that migrate in reducing environments. the concentrations of these elements exceed the quality standards (50 and 10 µg/l, respectively) in 22% and 1% of the samples. boron and selenium exceed the quality standards (300 and 10 µg/l, respectively) only in saline groundwater, where the concentrations of sulphate or chloride are significantly higher than quality standards for drinking water. in general, heavy metals in latvian groundwater are a smaller problem than in denmark where manganese and arsenic exceed quality standards of drinking water in 63% and 9% of the monitoring screens, respectively. however, these compounds do not pose a problem for the drinking-water supply in general in denmark due to sorption in natural sand filters at the watertreatment plants. in latvia, nickel concentrations higher than the quality standard of drinking water (20 µg/l) were not observed in groundwater below 10 m. in denmark, high nickel concentrations locally create problems where large-scale depression cones in the groundwater table result in oxidation of sulphide minerals. lowering of the groundwater table in latvian aquifers occurs only in deep, confined aquifers rather than in the vicinity of the groundwater table. conclusions the distribution patterns of minor elements in latvian and danish groundwater are somewhat similar and are governed by the basic groundwater chemistry. in general, the background concentrations of aluminium and fluoride are higher, and the concentrations of bromide, lithium, manganese, phosphorus and strontium are lower in shallow groundwater in latvia compared to the danish aquifers. the lower anthropogenic pressure combined with a higher confinement degree of major water-supply aquifers explains why diffuse contamination of latvian groundwater is lower than in denmark. at present, the limited data on pesticides and on the age of groundwater do not allow an assessment of the extent of the pesticide problem in latvian groundwater with sufficient accuracy. acknowledgements our much appreciated friend and colleague, edmund gosk, passed away on 24 november 2006 after a long struggle against cancer. we wish to acknowledge his engaged involvement in this project in latvia, as well as in his many other tasks. references geus 2005: grundvand 2004. status og udvikling 1989–2004. københavn: danmarks og grønlands geologiske undersøgelse (in danish with english summary). available on: www.geus.dk/publications/grund vandsovervaagning/1989_2004/index.html gosk, e., levins, i. & jørgensen, l.f. 2006: agricultural influence on groundwater in latvia. danmarks og grønlands geologiske under søgelse rapport 2006/85, 98 pp. henriksen, h.j. & sonnenborg, a. (eds) 2003: ferskvandets kredsløb. nova 2003 temarapport, 230 pp. københavn: danmarks og grøn lands geologiske undersøgelse, danmarks miljøundersøgelser, dan marks jordbrugsforskning and danmarks meteorologiske institut (in danish with english summary). also available on: http://vandmodel.dk/ferskvands_2003_final.htm neri 2005: novana – national monitoring and assessment programme for the aquatic and terrestrial environment. programme description, part 2. neri technical report 537/2005, 137 pp. stålnacke, p., grimvall, a., libiseller, c., laznik, m. & kokoite, i. 2003: trends in nutrient concentrations in latvian rivers and the response to the dramatic change in agriculture. journal of hydrology 283, 184– 205. stockmarr, j. 2005: groundwater quality monitoring in denmark. geo l og ical survey of denmark and greenland bulletin 7, 33–36. authors’ addresses l.f.j. & e.g. (deceased), geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: lfj@geus.dk i.l., latvian environment, geology and meteorology agency (legma), maskavas str. 165, lv-1019 riga, latvia. fig. 4. distribution of selected minor elements in shallow latvian groundwater (full lines indicate data from 477 wells and springs for the period 2003–2005) and in shallow danish groundwater (dotted lines indicate data from 723 monitoring screens for the period 1998–2003). geological survey of denmark and greenland bulletin 40 de nationale geologiske undersøgelser for danmark og grønland (geus) geological survey of denmark and greenland g eological survey of d enm ark and g reenland b ulletin 40 • 2018 volcanic rocks of the paleocene m aligât form ation, w est g reenland øster voldgade 10 dk-1350 copenhagen k denmark lithostratigraphy, geology and geochemistry of the volcanic rocks of the maligât formation and associated intrusions on disko and nuussuaq, paleocene of west greenland the upper cretaceous–tertiary nuussuaq basin in west greenland contains a many kilometres thick succession of siliciclastic sediments and overlying volcanic rocks. the first studies in the early 19th century were centred on the coal and fossils in the sediments and the minerals in the volcanic rocks, including famous occurrences of native iron. the present focus of interest includes modern stratigraphic and volcanological studies to decipher the basin evolution and support hydrocarbon and mineral exploration. this bulletin presents the lithostratigraphy, geology and geochemistry of the paleocene volcanic rocks of the maligât formation and its related intrusions on disko and the nuussuaq peninsula; it concludes with a detailed discussion of the effects of crustal contamination processes. the maligât formation is up to 2000 m thick and comprises four formally defined members and 15 chemically defined informal units. it is mainly composed of basalt lavas but also includes basaltic andesite, andesite and dacite flows and rhyolite tuffs. the silicic rocks and intrusions were produced by contamination in high-level magma chambers and commonly contain native iron. the comprehensive descriptions and analyses of each member and unit represent a synthesis of many years of work and are intended to serve as a guide for future studies, including exploration for mineral deposits associated with some units of the formation. a companion bulletin (volume 39) on the volcanic rocks of the picritic vaigat formation that underlies the maligât formation was published in 2017. lithostratigraphy lithostratigraphy, geology and geochemistry of the volcanic rocks of the maligât formation and associated intrusions on disko and nuussuaq, paleocene of west greenland asger ken pedersen, lotte melchior larsen & gunver krarup pedersen g e o l o g i c a l s u r v e y o f d e n m a r k a n d g r e e n l a n d b u l l e t i n 4 0 c o p e n h a g e n • 2 0 1 8 printed 2018 issn 1604-8156 isbn 978-87-7871-498-5 geus is a research and advisory institution in the danish ministry of energy, utilities and climate d e n m a r k geocenter denmark is a formalised cooperation between geological survey of denmark and greenland (geus), department of geoscience at aarhus university and the geological museum and department of geosciences and natural resource management at the university of copenhagen. 9 788778 714985 1 lithostratigraphy, geology and geochemistry of the volcanic rocks of the maligât formation and associated intrusions on disko and nuussuaq, paleocene of west greenland asger ken pedersen, lotte melchior larsen and gunver krarup pedersen geological survey of denmark and greenland ministry of energy, utilities and climate geological survey of denmark and greenland bulletin 40 • 2018 22 geological survey of denmark and greenland bulletin 40 keywords maligât formation, basalt, lithostratigraphy, intrusions, crustal contamination, native iron, graphite, nuussuaq basin, paleocene cover illustration uppalluk/giesecke monument, a landmark peak reaching 1574 m a.s.l. on the south coast of nuussuaq. subaerial lava flows of the upper rinks dal member of the maligât formation constitute the major part of the peak, which is topped by a remnant of a nordfjord member basalt flow. the exposed section is 550 m thick. photo: erik vest sørensen. frontispiece: facing page the north-east coast of disko in midnight sun. a wall of subaerial lava flows of the rinks dal member of the maligât formation rises above a slope dominated by numerous landslides. the slides are caused by an unstable configuration of basalts overlying poorly consolidated cretaceous to paleocene sediments. the ice-capped peak to the left is the 1366 m high qinngusaq mountain. in the foreground, the cutter porsild of the arctic station in qeqertarsuaq is anchored in the bay at tartunaq, nuussuaq, 1986. chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, department of geoscience, aarhus university; minik rosing, geological museum, university of copenhagen; finn surlyk, department of geosciences and natural resource management, university of copenhagen scientific editor: adam a. garde editorial secretary: jane holst referees: c.h. emeleus and r. wilson illustrations: jette halskov, willy l. weng, benny m. schark, jacob lind bendtsen and allan lindy. photographs: all photographs were taken by the authors except where otherwise stated. layout and graphic production: henrik klinge pedersen and jacob lind bendtsen printers: rosendahls-schultz grafisk a/s, albertslund, denmark manuscript received: 9 june 2017 final version approved: 15 may 2018 printed: september 2018 issn (print) 1604-8156 issn (online) 1904-4666 isbn (print) 978-87-7871-498-5 isbn (online) 978-87-7871-499-2 citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 40, 239 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark to buy bulletin in printed form please contact bogsalg@geus.dk and at www.geus.dk/bulletin40 (open access) © de nationale geologiske undersøgelser for danmark og grønland (geus), 2018 for the full text of the geus copyright clause, please refer to www.geus.dk/bulletin mailto:bogsalg@geus.dk http://www.geus.dk/bulletin40 http://www.geus.dk/bulletin 3 44 5 dedicated to the memory of finn ulff-møller 66 contents abstract .................................................................................................................................................................................................................................................................................................................. 9 introduction ................................................................................................................................................................................................................................................................................................... 10 geological setting ................................................................................................................................................................................................................................................................................. 14 the sedimentary substrate ....................................................................................................................................................................................................................................................... 16 deposition of the volcanic succession and equivalent sediments ......................................................................................................................................... 16 methods .............................................................................................................................................................................................................................................................................................................. 20 stratigraphic subdivision and numerical coding system ......................................................................................................................................................................... 20 nomenclature ........................................................................................................................................................................................................................................................................................... 20 geological sections and profiles .............................................................................................................................................................................................................................. 20 west greenland basalt group ............................................................................................................................................................................................................................................. 22 summary of the petrology of the vaigat and maligât formations ........................................................................................................................................ 22 maligât formation ................................................................................................................................................................................................................................................................................ 23 revision of the maligât formation ................................................................................................................................................................................................................... 23 lithostratigraphy of the maligât formation .................................................................................................................................................................................................... 23 general features of the maligât formation ........................................................................................................................................................................................................ 24 subdivision ........................................................................................................................................................................................................................................................................................... 24 lithological variations and the filling of the assoq lake ......................................................................................................................................................... 25 boundaries ............................................................................................................................................................................................................................................................................................. 39 lower boundary ...................................................................................................................................................................................................................................................................... 39 upper boundary ....................................................................................................................................................................................................................................................................... 42 rinks dal member ................................................................................................................................................................................................................................................................................ 43 lithostratigraphy of the rinks dal member ................................................................................................................................................................................................... 43 subdivision of the rinks dal member .................................................................................................................................................................................................................... 44 lower rinks dal member ....................................................................................................................................................................................................................................................... 45 summary of the main features of the lower rinks dal member .................................................................................................................................... 45 unit 505 (lowest flows) ....................................................................................................................................................................................................................................................... 46 unit 506 ................................................................................................................................................................................................................................................................................................... 46 unit 507 (low-ti flows) ...................................................................................................................................................................................................................................................... 47 unit 509 ................................................................................................................................................................................................................................................................................................... 55 south coast of disko: infilling of the southern part of the assoq lake .................................................................................................................... 56 the northern basin in the kuugannguaq–qullissat area .................................................................................................................................................... 61 unit 511 (skarvefjeld unit) ............................................................................................................................................................................................................................................ 65 south coast of disko around skarvefjeld .................................................................................................................................................................................................... 70 east disko around kvandalen .............................................................................................................................................................................................................................. 71 northt coast of disko around qullissat ....................................................................................................................................................................................................... 75 unit 512 (lower transition flows) ........................................................................................................................................................................................................................... 76 invasive lava flows on eastern disko ............................................................................................................................................................................................................... 77 invasive lava flows on southern nuussuaq .............................................................................................................................................................................................. 77 middle rinks dal member .................................................................................................................................................................................................................................................... 80 summary of the main features of the middle rinks dal member .............................................................................................................................. 80 unit 513 (akuarut unit) ..................................................................................................................................................................................................................................................... 80 south coast of disko between marraat qaqqaat and tuapaat qaqqaat ........................................................................................................... 83 south coast of nuussuaq ................................................................................................................................................................................................................................................ 85 upper rinks dal member ....................................................................................................................................................................................................................................................... 88 summary of the main features of the upper rinks dal member ................................................................................................................................. 88 unit 514 (upper transition flows) .......................................................................................................................................................................................................................... 88 units 515, 516 and 517 (the main upper rinks dal member) ........................................................................................................................................ 89 7 unit 518 (uppermost flows) ........................................................................................................................................................................................................................................... 91 lithologies of the upper rinks dal member ............................................................................................................................................................................................ 91 eruption sites on western disko ............................................................................................................................................................................................................................. 92 lithologies on disko ............................................................................................................................................................................................................................................................. 92 lithologies on nuussuaq ................................................................................................................................................................................................................................................ 99 lava volumes in the upper rinks dal member .................................................................................................................................................................................. 101 chemical compositions of the rinks dal member ............................................................................................................................................................................. 104 major elements ....................................................................................................................................................................................................................................................................... 104 trace elements ......................................................................................................................................................................................................................................................................... 104 eruption sites for the rinks dal member ........................................................................................................................................................................................................ 113 uncontaminated basalt dykes and sills ......................................................................................................................................................................................................... 113 feeder systems for the rinks dal member .............................................................................................................................................................................................. 114 nordfjord and niaqussat members ........................................................................................................................................................................................................................... 115 revision of the boundary between the nordfjord and niaqussat members .................................................................................................. 115 nordfjord member ............................................................................................................................................................................................................................................................................ 116 summary of the main features of the nordfjord member .................................................................................................................................................. 116 lithostratigraphy of the nordfjord member ............................................................................................................................................................................................... 116 internal structure of the nordfjord member ............................................................................................................................................................................................... 117 igneous rock types ........................................................................................................................................................................................................................................................................... 119 basalt ......................................................................................................................................................................................................................................................................................................... 119 basaltic andesite .......................................................................................................................................................................................................................................................................... 122 andesite ................................................................................................................................................................................................................................................................................................. 123 dacite ........................................................................................................................................................................................................................................................................................................ 124 rhyolite .................................................................................................................................................................................................................................................................................................. 128 geological themes and locality descriptions ................................................................................................................................................................................................. 131 sediment horizons at the base of the nordfjord member ................................................................................................................................................... 131 the west disko graphite rhyolite (wdgr) volcano ......................................................................................................................................................... 134 tuff successions ........................................................................................................................................................................................................................................................................ 136 conglomerates and volcaniclastic sandstones ..................................................................................................................................................................................... 138 sediments with clasts from the wdgr volcano ............................................................................................................................................................................ 141 sediments without clasts from the wdgr volcano .................................................................................................................................................................. 144 simple dacite lava flows between hammer dal and jamma .............................................................................................................................................. 144 composite lava flows ........................................................................................................................................................................................................................................................... 146 crater site lithologies ........................................................................................................................................................................................................................................................... 153 chemical compositions of the nordfjord member ............................................................................................................................................................................. 156 major elements ....................................................................................................................................................................................................................................................................... 156 trace elements ......................................................................................................................................................................................................................................................................... 161 composite lava flows ....................................................................................................................................................................................................................................................... 162 niaqussat member .............................................................................................................................................................................................................................................................................. 169 summary of the main features of the niaqussat member .......................................................................................................................................................... 169 lithostratigraphy of the niaqussat member ................................................................................................................................................................................................ 169 lower niaqussat member (unit 530) .................................................................................................................................................................................................................... 170 western disko .......................................................................................................................................................................................................................................................................... 171 eastern disko ........................................................................................................................................................................................................................................................................... 174 eastern nuussuaq .............................................................................................................................................................................................................................................................. 175 middle niaqussat member (unit 531) ................................................................................................................................................................................................................. 176 upper niaqussat member (unit 532) .................................................................................................................................................................................................................... 177 disko ................................................................................................................................................................................................................................................................................................... 178 eastern nuussuaq .............................................................................................................................................................................................................................................................. 179 chemical compositions of the niaqussat member .............................................................................................................................................................................. 182 major elements ...................................................................................................................................................................................................................................................................... 183 88 trace elements ........................................................................................................................................................................................................................................................................ 183 sapernuvik member ........................................................................................................................................................................................................................................................................ 192 summary of the main features of the sapernuvik member ............................................................................................................................................. 192 lithostratigraphy of the sapernuvik member .......................................................................................................................................................................................... 192 geology and geochemistry of the sapernuvik member ............................................................................................................................................................... 193 dyke systems of the nordfjord and niaqussat members .......................................................................................................................................................... 195 summary of the main features of the contaminated dyke systems ............................................................................................................................ 195 weakly contaminated dykes .............................................................................................................................................................................................................................................. 195 main dyke systems of the strongly contaminated magmas ........................................................................................................................................................ 195 intrusions in dyke system a .................................................................................................................................................................................................................................... 197 intrusions in dyke system b ..................................................................................................................................................................................................................................... 207 intrusions in dyke system c .................................................................................................................................................................................................................................... 208 other dyke intrusions .................................................................................................................................................................................................................................................... 209 concluding remarks on the strongly contaminated dyke systems .............................................................................................................................. 210 chemical compositions of the nordfjord and niaqussat member dykes and feeder systems .................................................... 210 volume relations of the maligât formation .................................................................................................................................................................................................. 215 thickness of the removed succession .............................................................................................................................................................................................................. 216 crustal contamination of the volcanic rocks .............................................................................................................................................................................................. 217 summary of the main features of the crustal contamination processes ............................................................................................................... 217 possible contaminants ........................................................................................................................................................................................................................................................ 217 basement ...................................................................................................................................................................................................................................................................................... 217 sediments ..................................................................................................................................................................................................................................................................................... 217 distinction between contaminants ................................................................................................................................................................................................................ 223 degrees of contamination ............................................................................................................................................................................................................................................. 224 contamination processes and geochemical changes during contamination .............................................................................................. 225 mixing and afc processes ...................................................................................................................................................................................................................................... 225 reduction processes ........................................................................................................................................................................................................................................................... 226 exchange of elements ..................................................................................................................................................................................................................................................... 228 concluding remarks ....................................................................................................................................................................................................................................................................... 230 acknowledgements ....................................................................................................................................................................................................................................................................... 230 references ...................................................................................................................................................................................................................................................................................................... 231 list of geological map sheets and sections ...................................................................................................................................................................................................... 236 appendix: place names .............................................................................................................................................................................................................................................................. 238 9 abstract pedersen, a.k., larsen, l.m. & pedersen, g.k. 2018: lithostratigraphy, geology and geochemistry of the volcanic rocks of the maligât formation and associated intrusions on disko and nuussuaq, paleocene of west greenland. geological survey of denmark and greenland bulletin 40, 239 pp. the paleocene volcanic rocks in the nuussuaq basin on disko and nuussuaq comprise the picritic vaigat formation (c. 62–61 ma) and the overlying basaltic maligât formation (c. 60 ma). the maligât formation is up to 2000 m thick on western disko where the top of the formation is least eroded. the formation is divided into four members, the rinks dal, nordfjord, niaqussat and sapernuvik members, which are formally defined here. on central and eastern disko and nuussuaq the maligât formation lavas are interbedded with fluvial and lacustrine sandstones and mudstones of the atanikerluk formation. the rinks dal member is the lowest member and originally constituted around 61% by volume of the formation. it is divided into 12 informal units based on chemically recognisable oscillations in the fractionation state of the basalts. the oldest units are present on central and south disko close to the disko gneiss ridge. the younger lavas spread farther to the east, north and west, filled the assoq lake basin east of the ridge and gradually onlapped the shield of the earlier vaigat formation that rose to the north. only the lavas of the upper rinks dal member reached far into nuussuaq. the lavas are generally not crustally contaminated and comprise evolved basalts with 4.4–9.2 wt% mgo and a few picrites. the most evolved basalts with 3.2–4.8 wt% tio2 occur in the middle part of the member where they form the akuarut unit. the nordfjord member originally constituted around 6% by volume of the formation. it is not subdivided because the lithological variability is local. the member is widespread but has its depocentre on north-western disko where thicknesses reach 350 m and eruption sites, intermediate lavas and acid tuffs are present. over most of the area the member consists of just a few lava flows with combined thicknesses of 30−100 m. the member has a very diverse lithology with rock types ranging from silicic basalt with 5.3–10.0 wt% mgo through magnesian basaltic andesite and andesite with 2.4–10.6 wt% mgo to dacite with 1.2–2.2 wt% mgo. rhyolite with 0.2–1.2 wt% mgo and up to 77 wt% sio2 occur in tuffs and conglomerate clasts. all rocks are crustally contaminated and some are native-iron-bearing. the niaqussat member originally constituted around 33% by volume of the formation. it is subdivided into three informal units. the member is widespread, but much of it has been removed by erosion. lithologies in the lower unit range from silicic picrite with up to 15 wt% mgo to basalt with 6–12 wt% mgo and a few basaltic andesite flows. the middle and upper parts of the niaqussat member comprise more evolved basalts with respectively 6.1–7.2 wt% mgo and 4.9–6.4 wt% mgo. all rocks are crustally contaminated and a few lava flows are native-iron-bearing. the sapernuvik member comprises three uncontaminated basalt flows with 7.5–10.7 wt% mgo. it is only preserved in a small area on western disko. dyke systems with up to 80 km long dykes and subvolcanic intrusions associated with the nordfjord and niaqussat members occur on western and north-eastern disko. the rocks are crustally contaminated and range from silicic basalt with 4–13 wt% mgo to magnesian andesite with 3–10 wt% mgo. they commonly form composite intrusions, some of which contain accumulations of native iron and sulfides. the contaminants are carbonand sulfur-bearing sediments of the nuussuaq group. major contamination mechanisms were mixing with partial melts from the sediment sidewall and xenoliths and selective exchange of some elements, including carbon and sulfur, between magma and sediment. degrees of contamination vary from 2−5% in the basalts to 10−50% in the more silicic rocks. no rocks more evolved than basalt were produced by ordinary fractional crystallisation. a.k.p., natural history museum of denmark, university of copenhagen, øster voldgade 5-7, dk-1350 copenhagen k, denmark, and geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: akp@snm.ku.dk l.m.l., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: lml@geus.dk g.k.p., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: gkp@geus.dk 1010 introduction the nuussuaq basin is one of a series of linked basins that extend along the entire western continental margin of greenland with a basin fill comprising mesozoic to tertiary sediments and tertiary volcanic rocks (chalmers & pulvertaft 2001). the basins are partly situated in the offshore areas, but between 69° and 73°n the rocks of the nuussuaq basin are exposed on disko, nuussuaq, ubekendt ejland and svartenhuk halvø (fig. 1). the nuussuaq basin contains a cretaceous to paleocene sedimentary succession overlain by a paleocene to eocene volcanic succession. the rocks have been studied since the mid-1850s with an academic focus on the wellpreserved fossils and the occurrence of native iron in the volcanic rocks. the economic interest has been focused on the coal beds and, in later years, on the presence of hydrocarbons. the lithostratigraphy of the sediments (the nuussuaq group) was defined and described by dam et nn nuussuaq 71°n 72°n 72°n 51°w53°w 71°n 70°n 69°n 55°w 51°w svartenhuk halvø ubekendt ejland greenland ice sheet uummannaq innerit hareøen itil li fau lt vaigat disko disko bugt ilulissat qeqertarsuaq aasiaat ebf ebf 50 km greenland neogene sediment cover offshore saqqaata qaqqaa central complex paleocene picrites (vaigat formation) undifferentiated basalts offshore naqerloq formation svartenhuk formation maligât formation maastrichtian– paleocene sediments albian–campanian sediments extensional fault precambrian basement fault with lateral or alternating displacements gb01_02_019_03_lml.eps fig. 1. simplified map of the pre-quaternary geology of the nuussuaq basin. light colours denote sea-covered areas. ebf: eastern boundary fault system. 11 al. (2009). the present work on the volcanic rocks may be considered a sequel to this work. the volcanic rocks in the nuussuaq basin crop out over c. 20  000 km2 within a basinal area of c. 120 km × 400 km (c. 50 000 km2); however, volcanic rocks extend over a total area of c. 200 km × 550 km (c. 110 000 km2), much of which is below sea level (chalmers et al. 1999; oakey & chalmers 2012). the rocks in the nuussuaq basin were erupted at highly variable rates into a tectonically very active environment, and there was a complex interplay between the volcanic, sedimentary and tectonic evolution of the basin, which resulted in a highly complex architecture of the volcanic succession. moreover, the range of magma compositions was large. west greenland contains an unusually high proportion of primitive, mg-rich picritic rocks that represent almost unmodified mantle melts (drever 1953, 1956; clarke 1970; clarke & upton 1971; clarke & pedersen 1976; ? ? ? nûluk mb ordlingassoq mb b e l o w e x p o s u r e l e v e l ordlingassoq mb qeqertalik mb tuperssuartâta kûa mb nunavik mb nerutusoq mb kakilisaat mb 60.2 ± 0.5 niaqussat mb dykes 57.49 ± 1.40 58.34 ± 0.40 dykes 57.49 ± 1.40 58.34 ± 0.40 nordfjord mb rinks dal mb 61.2 ± 0.4 ordlingassoq mb 61.3 ± 0.5 naujánguit mb 61.2 ± 0.5 anaanaa mb 62 (c27n) 58.66 ± 0.34 54.03 ± 0.33 48.02 ± 2.76 56.15 ± 0.41 58.31 ± 0.30 54.86 ± 0.32 alkaline sill 54.52 ± 0.67 dykes 54.3 ± 0.3 55.5 ± 0.4 56.8 ± 0.8 alkaline sill 54.52 ± 0.67 dykes 54.3 ± 0.3 55.5 ± 0.4 56.8 ± 0.8 gabbro sill 58.96 ± 0.51 gabbro sill 58.96 ± 0.51 avatarpaat plug 27.8 ± 0.6 h ar eø en fo rm at io n n aq er lo q fo rm at io n s va rte nh uk fo rm at io n m al ig ât fo rm at io n va ig at fo rm at io n e rq ua fm 39 –3 8 56 –5 4 60 –5 8 61 –6 0 e oc en e p al eo ce ne 62 –6 1 54 –5 3 talerua mb 38.74 ± 0.23 aumarûtigssâ mb dykes dykes 34.5 ± 0.2 ubekendt ejlanddisko nuussuaq ma hareøen and west nuussuaq west of itilli fault east of itilli fault svartenhuk halvø erqua fm 53.47 ± 0.52 naqerloq fm 55.91 ± 0.60 54.86 ± 0.44 arf. trachyte 57.51 ± 0.24 skalø mb 57.98 ± 0.59 nuuit mb 58.05 ± 0.59 tunuarsuk mb 59.41 ± 0.61 sediments 55.94 ± 0.20 u. m. l. u. m. l. s.q. gv02_02_047_lml 57.25 ± 0.95 56.99 ± 0.49 55.21 ± 0.30 59.50 ± 1.71 59.97 ± 0.89 60.31 ± 1.39 kanísut mb nûk takisôq mb ifsorisok mb c 27 n c 26 r c 27 r c 26 r c 26 r c 24 r c 24 r c 18 n c 17 r fig. 2. stratigraphic scheme for the volcanic rocks in the nuussuaq basin, based on hald & pedersen (1975) for disko and nuussuaq, hald (1976) for hareøen and western nuussuaq, j.g. larsen (1977) for ubekendt ejland, and j.g. larsen & grocott (1991) and j.g. larsen & pulvertaft (2000) for svartenhuk halvø. radiometric ages (ma) with one digit after the decimal point are from storey et al. (1998) and larsen et al. (2009); ages with two digits after the decimal point are from larsen et al. (2016). the age for the anaanaa member is not radiometric but based on its normally magnetised character. note that the vertical ‘age scale’ is not equidistant. s.q.: saqqaata qaqqaa central complex. l., m. and u. denote lower, middle and upper. arf: arfertuarsuk trachyte. sediments with a quartzo-feldspathic component are yellow; purely volcaniclastic sediments are brown. wavy lines indicate unconformities. palaeomagnetic directions with magnetochrons are indicated at the right side of some lithological units, from riisager & abrahamsen (1999), riisager et al. (1999, 2003), schmidt et al. (2005) and unpublished data by p. riisager, 2006). modified from larsen et al. (2016). 1212 holm et al. 1993; larsen & pedersen 2000, 2009). however, repeated episodes of assimilation of crustal material, including organic-rich sediments, led to formation of units of silica-enriched rocks, including native-ironand graphite-bearing basalts, andesites and dacites as well as graphite-bearing rhyolites (e.g. törnebohm 1878; steenstrup 1883; pedersen 1977a, 1981, 1985; pedersen & larsen 2006; larsen & pedersen 2009). the volcanic succession is divided into two major parts. the lower part is dominated by picrites and has been formalised as the vaigat formation, which is present from disko and northwards to svartenhuk halvø (hald & pedersen 1975; clarke & pedersen 1976). the upper part is dominated by basalts and has previously been referred to different formations in different areas; based on 40ar–39ar age determinations larsen et al. (2016) concluded that the basalts on disko and nuussuaq east of the itilli fault belong to the maligât formation which is not present in other areas (figs 1, 2). this bulletin on the maligât formation and its companion on the vaigat formation (pedersen et al. 2017) describe the lithostratigraphy, geology and geochemistry of the volcanic succession on disko and nuussuaq east of the itilli fault. they represent a synthesis of work that has been going on since 1968 and from which many results have been published. the two bulletins may be viewed as a comprehensive map description covering the paleocene volcanic rocks in all the geological maps of the area, which is covered by seven geological maps on a scale of 1:100 000, two geological maps on scales of 1:50 000 and 1:20 000 covering different parts of nuussuaq, and five vertical geological sections on a scale of 1:20 000 across disko and nuussuaq (fig. 3). an important task for the two bulletins has been the definition and documentation of the several volcanic units that appear on the published maps and sections but have not been described before. the documentation includes geochemical data because chemostratigraphy is important in both formations and crucial for the maligât formation. emphasis is placed on descriptions with pictures of the variable lithologies in selected key areas and on geochemical descriptions and plots of the stratigraphic units, facilitated by a numerical coding system, data tables and a comprehensive electronic appendix with analytical data. in short, the bulletins are intended to serve as guides for future field, geochemical and petrological studies. the companion bulletin on the vaigat formation (pedersen et al. 2017) contains detailed introductory chapters on previous investigations, geological setting and methods that also cover the maligât formation; however short chapters on the geological setting and methods are also included here because they are necessary for the independent use of the present bulletin. petrogenetic considerations have been published by larsen & pedersen (2000, 2009) and larsen et al. (2003), and the two bullegv01_02_160_01_lml 70°n 69°30'n 55°w 53°w 55°w 53°w 25 km 70°30'n 70 v.1 nord agatdal 70 v.1 syd qutdligssat 69 v.2 nord pingu 69 v.1 syd uiffaq 69 v.1 nord mellemfjord 70 v.2 nord agpat special map paatuut fig. 103 fig. 173 ice cap west greenland basalt group nuussuaq group precambrian basement map sheets 1:100 000 special map 1:50 000 sections and map 1:20 000 maps in this volume fig. 3. index map of published geological maps and photogram metric sections covering disko and nuussuaq. also shown are the local geological maps that appear as figures in this bulletin. a list of all published maps and sections is found at the end of the reference list. 13 tins therefore only give short summaries of the petrology of the rocks. on the other hand, the present bulletin contains a final chapter on crustal contamination of the volcanic rocks, with descriptions of possible contaminants and discussion of contamination processes, that covers both the vaigat and maligât formations. our late friend and colleague finn ulff-møller studied the native-iron-bearing intrusions on western disko extensively, but at his death in 2009, much of his knowledge was left unpublished. in this bulletin, the chapter on dyke systems of the nordfjord and niaqussat members includes much information extracted from his phd thesis, field notes, sketches, photographs and samples. 1414 geological setting the tertiary volcanic rocks in west greenland form part of the north atlantic igneous province. the volcanism was associated with the breakup of the north atlantic craton, as described by noe-nygaard (1974), upton (1988), saunders et al. (1997) and pedersen et al. (2017). the c. 20 000 km2 of exposed volcanic rocks in west greenland make it the second largest exposed volcanic succession on the continental margins of the north atlantic, surpassed only by the plateau basalts in central east greenland. the exposed rocks in the nuussuaq basin consist of cretaceous to palaeogene sediments and palaeogene volcanic rocks. the three most important structural elements of the nuussuaq basin on disko and nuussuaq are the eastern boundary fault system (ebf), the kuugannguaq−qunnilik (k−q) fault system and the itilli fault system (figs 1, 4). these fault systems divide the basin into a central−eastern part where the volcanic rocks are close to flat-lying and their substrate is well exposed, a western part where the substrate is rarely exposed and the volcanic succession is faulted and commonly tilted, and a westernmost part west of the itilli fault system where, due to kilometre-scale downthrow to the north-west, a very different and mainly younger succession is preserved. the eastern boundary fault system delimits the basin towards the elevated precambrian basement in the east. the system consists of a number of fault segments with orientations from n−s to nw−se. within the basin, important faults run mainly n−s. the disko gneiss ridge is a pre-basaltic fault block that is down-tilted towards the east, whereas to the west it is delimited by a n−s-running fault system that is an extension of the k–q fault system. the surface of the gneiss ridge is shaped by prevolcanic erosion (bonow 2005); the later lavas have lapped onto the hilly landscape of the ridge and eventually covered it. the ridge stands up to c. 700 m a.s.l. on central disko, lowers to c. 300 m a.s.l. on southern disko and extends southward, mainly below sea level, across disko bugt to the low-lying gneiss areas in the aasiaat area. the ridge is not exposed at the north coast of disko; this has been attributed to a nw−se-trending fault on northern disko (chalmers et al. 1999). the k−q fault system runs s−n across both disko and nuussuaq, and pre-, synand postvolcanic movements have taken place along the faults. a clustering of volcanic eruption sites along the fault system demonstrates its deep-seated nature as a preferred pathway for magmas. west of the k–q fault system no prevolcanic sediments are exposed on disko, but sediment-contaminated volcanic rock units and sandstone and mudstone xenoliths in these attest to the presence of sediments at depth (e.g. pedersen 1977a). thick sedimentary successions are known from nuussuaq west of the k–q fault system (dam et al. 2009). western disko, west of the k–q fault system, is cut by several n−s-oriented post-paleocene faults with steeply e-dipping fault planes. they divide the volcanic succession into a number of antithetic fault blocks in which the lavas dip 2−22° w, thus exposing the youngest parts of the succession in the westernmost blocks. the antithetic fault system does not continue offshore but rather has given rise to a monocline of limited width (skaarup 2002; skaarup & pulvertaft 2007; gregersen & bidstrup 2008). according to refraction seismic data, paleocene oceanic crust is situated 120 km west of disko (funck et al. 2012; oakey & chalmers 2012). the itilli fault system is a major structural boundary (fig. 1; chalmers & pulvertaft 2001). downthrow is to the nw, with net displacements of more than 1 km on hareøen and nuussuaq. as a result, the volcanic successions on the two sides of the itilli fault system are fundamentally different (fig. 2). some relatively young (latest paleocene to eocene) volcanic formations occur west of the fault system but not east of it; these have not been included in this work. facing page: fig. 4. simplified geological map of disko and nuussuaq with locations of the major profiles used for the mapping of the maligât formation (blue dots) and also shown in the stratigraphic panels of figs 7–14. a few profiles are not shown. profiles in the vaigat formation (red dots) are shown for completeness. some localities are measured points with altitudes as given on the 1: 250  000 scale topographic maps. abbreviated names: ak: akuliarusersuaq. dj dal: daugaardjensen dal. eq: eqalunnguaqqat qaqqaat. fl dal: frederik lange dal. gd: giesecke dal. iko: ikorfarsuit. iq: ivissussat qaqqaat. kin: kingittuusaq. or: orlingasoq. qas.: qasigissat. si: saqqarliit ilorliit. ebf: eastern boundary fault system. k–q fault: kuugannguaq–qunnilik fault. additional localities are shown in fig. 6 and for north-west disko in fig. 103. note that on this map glaciers are not shown; the extent of the nordfjord member on central eastern disko, which is largely covered by the sermersuaq ice cap, is interpolated between many small exposures at the margin of the ice cap. 15 eocene lavas and intrusions nordfjord and niaqussat mbs rinks dal member maligât formation nordfjord vaigat qeqertarsuaq / godhavn innarsuaq / skarvefjeld vaigat formation st or da l cretaceous–tertiary sediments precambrian basement gv01_02_136_01_lml_03 fault profile in the vaigat (red) and maligât (blue) formations 25 km sapernuvik member kuugannguaq malig aat marraat anaanaa sikillingi nuusaq hareøen niaqussat d i s k o g n e i s s r i d g e asuk asuutaa nuuk killeq ataata kuua mellemfjord sapernuvik qas. aaffarsuaq ikorfat saqqaq niuluut killiit / fortunebay uiffaq tunup qaqqaa eqaluit perlertut qaqqaat pyramiden assoq skansen tuapaat qaqqaat marraat qaqqaat pingu skorstensfjeld aqajaruata qaqqaa qullissat inussuk paatuut niiortuut nunavik eqi iq or gd eqip qaqqaa giesecke monument umiusat atanikerluk tartunaq narsap qaqqaa qinngusaq sorte hak d. j. dal ak eq si iko kin orpiit qaqqaat kvandalen fl dal uummannaqkuuk point 2080 m 1640 m point 1530 m point 1510 m 1578 m 1012 m 1300 m inngigissoq 1123 m 1014 m 1888 m qaarsut kq fa ul t kq fa ul t itilli fault 69°30’n 70°30’n 70°n 55°w 53°w 52°w nuussuaq disko kangerluk / disko fjord ebf nûluk member nordfjord laksedalen brededal blåbærdalen rink dal hammer dal vesterdalen sortebærdalen 1616 the sedimentary substrate the nuussuaq basin was a depositional area for clastic sediments for at least 40 ma from the albian to the paleocene (dam et al. 2009). geophysical investigations have shown sediment thicknesses of at least 6 km and possibly up to 10 km in the western part of the basin (christiansen et al. 1995; chalmers et al. 1999); in the eastern part the sediments are only 1−2 km thick. most of the exposed sediments were deposited in a large delta complex which fanned out from south-east of disko towards the west and north-west; thus fluvial to deltaic environments predominate on disko and southern nuussuaq, while marine deep-water environments are found on north-western nuussuaq and on svartenhuk halvø (g.k. pedersen & pulvertaft 1992). the albian to campanian non-marine and marginally marine sediments are referred to the kome, slibestensfjeldet, upernivik næs and atane formations (dam et al. 2009) and comprise yellowish quartzo-feldspathic sandstones interbedded with mudstones and coal layers (midtgaard 1996; g.k. pedersen et al. 2006). dark marine mudstones of the itilli and kangilia formations are exposed on north-western nuussuaq. on western nuussuaq, the itilli formation is more than 2 km thick in the gro#3 well. the itilli formation comprises upper cretaceous turbiditic sandstones and mudstones deposited on a submarine slope west of the kuugannguaq−qunnilik fault (dam & sønderholm 1994; dam et al. 2009). an early campanian angular unconformity is observed on central nuussuaq (dam et al. 2000) and interpreted as representing a regional event (gregersen et al. 2013). during the maastrichtian and early paleocene, the basin was affected by faulting and changes in relative sea level. at least three phases of uplift during the late cretaceous and early paleocene are interpreted from deeply incised submarine canyons with conglomerates, turbidites and mudstones, as well as an incised fluvial valley with pebbly sandstones as the dominant lithology (dam & sønderholm 1994, 1998; dam 2002; dam et al. 2009). deposition of the volcanic succession and equivalent sediments at the onset of the volcanism, a marine embayment stretched from the open sea in the north-west (western nuussuaq and beyond this) towards south and south-east where it was bounded by the disko gneiss ridge in the west, an active delta depositing sediments in the south and south-east, and the gneiss highland in the east and north-east (dam et al. 2009, fig. 12). detailed studies of volcanic rocks sediments m al ig ât f o rm at io n va ig at f o rm at io n tunoqqu mb qo. mb asuk mb manîtdlat mb skarvefjeld unit akuarut unit lower rinks dal member upper rinks dal member ri nk s d al m em be r n iaq us sa t m em be r an aa na a m em be r abraham mb pingu mb umiussat mb assoq mb tuffs ? ? sapernuvik member upper middle lower nordfjord member st uk ni he tb tanuuk killeq mb nus. qa. mb marraat unit niaqornaq unit picrites and basalts, 5 units kûgánguaq mb at a n ik er lu k f o rm at io n eq a lu li k f o rm at io n ap cbs gv02_02_039_lml naujánguit member ordlingassoq member akunneq mb n au jât m b fig. 5. stratigraphic subdivision of the paleocene volcanic rocks and sediments on disko and nuussuaq. the ‘thicknesses’ of the volcanic units in the diagram are intended to approximately reflect their volumes, although this is not always possible. the units with elliptical outlines do not extend laterally throughout the contemporaneous succession, and the small ones are distinctly local. crustally contaminated units are shown in red, alkaline units in purple (with a red outline if crustally contaminated), and geochemically enriched units have dashed outlines. abbreviations (in stratigraphic order): nus. qa. mb: nuusap qaqqarsua member. he: henderson unit. ta: tunorsuaq-a unit. ni: niiortuut unit. tb: tunorsuaq-b unit. uk: ukallit unit. st: stordal alkaline unit. qo. mb: qordlortorssuaq member. cbs: contaminated basalts in stordal. ap: alkali picrites associated with the manîtdlat member. the lower and upper rinks dal member are further subdivided into a number of chemostratigraphic units not shown here (but see fig. 21). sediment members from dam et al. (2009): marine mudstones are grey, non-marine mudstones are brownish-grey, sandstones are yellow and coal layers are black. the tuffs correlated with the nordfjord member were found in 2016. 17 volcanic facies transitions have documented very high subsidence rates in excess of 20 m/ka during the deposition of the early part of the vaigat formation (pedersen et al. 2002b), making regional palaeogeographical reconstructions difficult. deposition of the volcanic rocks began in the north-west and took place concomitantly with continued deposition of sediments in other parts of the basin, leading to a complex interplay between volcanic and sedimentary members. a lithostratigraphic scheme for both volcanic rocks and sediments is shown in fig. 5. lateral basin fill from west to east, shifting eruption sites with time and frequent facies changes between subaerial and subaqueous facies within most volcanic units resulted in an unusually complex stratigraphy of the volcanic succession, as described in pedersen et al. (2017). in the volcanic succession, the vaigat formation is dominated by primitive, olivine-rich picritic rocks, which in subaerial facies form thin, grey, crumbly lava flows. the overlying maligât formation is dominated by evolved, plagioclase-phyric to aphyric basalts which in subaerial facies form thicker, darker grey to brownish, massive lava flows. both formations contain a number of lithologically and chemically distinct units; several of these form marker horizons and have been assigned member status (fig. 5). the vaigat formation forms a s–n elongate, 260 km long, shield-like feature in the nuussuaq basin. it is thickest, up to c. 5 km, on ubekendt ejland north of nuussuaq ( j.g. larsen 1977). on central and western nuussuaq it is up to 1.6 km thick, and on northern disko up to 1.5 km; it wedges out eastwards on eastern nuussuaq and eastwards and southwards on eastern and north-central disko (fig. 1). its lower boundary towards the sediments is strongly diachronous and youngs towards the east because of progradation of the volcanic pile from west to east. it was deposited in three major episodes of activity, which gave rise to its three main members (anaanaa, naujánguit and ordlingassoq). the marine embayment with mudstones of the eqalulik formation (dam et al. 2009) was gradually filled from the west, and at the end of the second episode the advancing volcanic front reached the eastern gneiss highland on north-east nuussuaq and blocked the connection to the sea; as a result the embayment was transformed into a freshwater lake, the naajaat lake (piasecki et al. 1992; pedersen et al. 1996), in which sediments of the non-marine atanikerluk formation were deposited (koch 1959; dam et al. 2009). this lake persisted through the third episode of the vaigat formation with deposition of mudstones of the naujât and pingu members, but it was greatly diminished towards the end of the episode by the combined infill of volcanic rocks and sandstones of the umiussat member (g.k. pedersen et al. 1998; dam et al. 2009). the maligât formation is centred on disko, and its oldest exposed part is found on south-central disko. it is thickest on western and central disko where its base is generally below exposure level; on eastern disko it rests on sediments, and in the north it onlaps the vaigat formation with increasingly younger lavas at the contact going north from disko to nuussuaq. eastwards the maligât formation thins and interfingers with lake sediments of the assoq lake basin and the sand-dominated delta deposits from a large, north-westto west-flowing river (dam et al. 2009). as a result, the eastern fringes of the formation changed with time from subaqueous lava flows and hyaloclastites through invasive flows intruded like sills into the wet sediments, to subaerial lava flows, which finally covered much of the basinal area and stepped onto the gneiss highland on eastern nuussuaq. both the vaigat and maligât formations are paleocene (danian–selandian) in age. the oldest part of the vaigat formation is normally magnetised and was emplaced during geomagnetic chron c27n (riisager & abrahamsen 1999), 62.5–62.2 ma (vandenberghe et al. 2012). the middle and upper parts of the vaigat formation, and all of the maligât formation, are reversely magnetised (deutsch & kristjansson 1974; athavale & sharma 1975; riisager et al. 1999), fitting into the long-lived geomagnetic chron c26r. the top of the maligât formation has an 40ar–39ar age of 60.2 ± 0.5 ma (storey et al. 1998). in the northern part of the nuussuaq basin, the basalts are referred to the svartenhuk, naqerloq and erqua formations (fig. 2; j.g. larsen 1977; j.g. larsen & grocott 1991; larsen et al. 2016). these are younger than the maligât formation except for the lowest part of the svartenhuk formation, which may be coeval with it. the 60–58 ma svartenhuk formation is present on svartenhuk halvø and north of it, on ubekendt ejland, and possibly also on hareøen and on nuussuaq west of the itilli fault (larsen et al. 2016; figs 1, 2). it is overlain by the 56–54 ma eocene naqerloq formation in the same areas, followed by the 54–53 ma erqua formation on ubekendt ejland. dykes similar to the basalts of the svartenhuk and naqerloq formations are widespread on disko and nuussuaq (larsen et al. 2016). a map of disko and nuussuaq with detailed place names used in this bulletin is shown in fig. 6. 1818 n iaq or na t iti lli iti lli m ar ra at k illi it h ar eø en m al ig aa t vai gat jam m a m an iill at as uk ki lle rp aa t q aq qa rs ua t q ul lis sa t in us su k g am le q ul lis sa t n un ng ar ut at an ike rlu k ta rt un aq ke gle n 55 °w 54 °w 53 °w 52 °w 70 °n 70 °3 0' n py ra m id en am et ys ts kr æn te n ag at fje ld et steenstrup dal as uu ta a o rli ng as oq ha ra ld mo ltk e d al ku ug an ng ua q kuugannguaq h am m er d al ri nk d al pa at uu t n iio rt uu t aa ffa rs ua q po in t 1 72 2 m n un av ik saqqaqdalen eq i ki ng itt oq sa qq aq g ie se ck e m on um en t agat dal en at aa ta ku ua qunnilik ik or fat q aa rs ut u um m an na q po in t 2 08 0 m aa ffa rs ua q po in t 8 82 m po in t 5 00 m po in t 4 40 m n iaq us sa t se di m en tk lø fte n po in t 7 80 m po in t 1 07 0 m h an ek am m en po in t 1 65 0 m po in t 1 38 0 m po in t 1 51 0 m po in t 1 46 0 m po in t 1 56 0 m po in t 1 61 0 m po in t 1 59 0 m ill ul ua rs ui t q aq qa a gi es ec ke d al morte n p or sil d dal po in t 2 00 0 m m ak itt ar iss ag aqu m iar to rfi up q aq qa a po in t 1 88 8 m iv iss us sa t q aq qa at po in t 2 01 0 m po in t 1 76 0 m eq ip q aq qa a u m iu sa t po in t 9 75 m ’sa qq aq ’ pu ia ttu ssu aq 0 10 20 km 69 °3 0' n stordal sto rda l ka ng er so oq / no rd fjo rd ak ull iit / m ell em fjo rd g v0 1_ 04 _1 9_ 03 _l m l u un ar tu ar su k ill uk un ng ua q in ng igi ss oq ak un ne q pi ng u aq aja ru at a q aq qa a sk an se n tu ap aa t tu ap aa t q aq qa at sin iff ik m ar ra at q aq qa at n iu lu ut as so q ip pi k in na rs ua q / s ka rv ef je ld q eq er ta rs ua q / g od ha vn ak ua ru t / l yn gm ar ks fje ld tu ap as su itti ni n av ar an aa t / a po st el fje ld ki llii t / f or tu ne ba y lu cie fje ld n uu k ka ng ille q ka ng er luk / di sk o fjo rd rødeelv blæsedalen brededal sk or st en sfj el d laksedalen blå bæ rda len po in t 1 02 5 m po in t 1 01 4 m po in t 1 12 3 m s o rt ebærdalen kvan dal en fr ed er ik l an ge d al c ha rle s p ol ar is d al dau ga ar djen se n d al ak ul iar us er su aq kuannersuit sullua t eq alu nn gu aq qa t q aq qa at po in t 1 10 9 m o rp iit q aq qa at kuannersuit kuussuat so rt e h ak po in t 1 44 0 m po in t 1 26 6 m ill uk as ik ku an ni t n aq er lo q tu nu p q aq qa a u ka le qa rt ar fik eq alu it/ n or dr e la ks eb ug t sa pe rn uv ikki ng itt up q aq qa a iv isa ar qu t / en ok h av n sa qq ar liit il or liit sa qq ar liit s ila rli it po in t 1 01 2 m pe rle rt ut q aq qa at po in t 1 30 0 m po in t 1 13 7 m po in t 1 53 0 m n ar sa p q aq qa a q in ng us aq po in t 1 57 8 m ik or far su it je rn py nt enqa sig iss at ki ng itt uu sa q ve st er da len po in t 1 13 2 m sermersuaq ser mers uaq po in t 1 64 0 m ku ug an ng ua q se kv an da le n u iff aqal an ng up q aq qa i n iaq or na t iti lli iti lli m ar ra at k illi it h ar eø en m al ig aa t vai gat jam m a m an iill at as uk ki lle rp aa t q aq qa rs ua t q ul lis sa t in us su k g am le q ul lis sa t n un ng ar ut at an ike rlu k ta rt un aq ke gle n 55 °w 54 °w 53 °w 52 °w 70 °n 70 °3 0' n py ra m id en am et ys ts kr æn te n ag at fje ld et steenstrup dal as uu ta a o rli ng as oq ha ra ld mo ltk e d al ku ug an ng ua q kuugannguaq h am m er d al ri nk d al pa at uu t n iio rt uu t aa ffa rs ua q po in t 1 72 2 m n un av ik saqqaqdalen eq i ki ng itt oq sa qq aq g ie se ck e m on um en t agat dal en at aa ta ku ua qunnilik ik or fat q aa rs ut u um m an na q po in t 2 08 0 m aa ffa rs ua q po in t 8 82 m po in t 5 00 m po in t 4 40 m n iaq us sa t se di m en tk lø fte n po in t 7 80 m po in t 1 07 0 m h an ek am m en po in t 1 65 0 m po in t 1 38 0 m po in t 1 51 0 m po in t 1 46 0 m po in t 1 56 0 m po in t 1 61 0 m po in t 1 59 0 m ill ul ua rs ui t q aq qa a gi es ec ke d al morte n p or sil d dal po in t 2 00 0 m m ak itt ar iss ag aqu m iar to rfi up q aq qa a po in t 1 88 8 m iv iss us sa t q aq qa at po in t 2 01 0 m po in t 1 76 0 m eq ip q aq qa a u m iu sa t po in t 9 75 m ’sa qq aq ’ pu ia ttu ssu aq 0 10 20 km 69 °3 0' n stordal sto rda l ka ng er so oq / no rd fjo rd ak ull iit / m ell em fjo rd g v0 1_ 04 _1 9_ 03 _l m l u un ar tu ar su k ill uk un ng ua q in ng igi ss oq ak un ne q pi ng u aq aja ru at a q aq qa a sk an se n tu ap aa t tu ap aa t q aq qa at sin iff ik m ar ra at q aq qa at n iu lu ut as so q ip pi k in na rs ua q / s ka rv ef je ld q eq er ta rs ua q / g od ha vn ak ua ru t / l yn gm ar ks fje ld tu ap as su itti ni n av ar an aa t / a po st el fje ld ki llii t / f or tu ne ba y lu cie fje ld n uu k ka ng ille q ka ng er luk / di sk o fjo rd rødeelv blæsedalen brededal sk or st en sfj el d laksedalen blå bæ rda len po in t 1 02 5 m po in t 1 01 4 m po in t 1 12 3 m s o rt ebærdalen kvan dal en fr ed er ik l an ge d al c ha rle s p ol ar is d al dau ga ar djen se n d al ak ul iar us er su aq kuannersuit sullua t eq alu nn gu aq qa t q aq qa at po in t 1 10 9 m o rp iit q aq qa at kuannersuit kuussuat so rt e h ak po in t 1 44 0 m po in t 1 26 6 m ill uk as ik ku an ni t n aq er lo q tu nu p q aq qa a u ka le qa rt ar fik eq alu it/ n or dr e la ks eb ug t sa pe rn uv ikki ng itt up q aq qa a iv isa ar qu t / en ok h av n sa qq ar liit il or liit sa qq ar liit s ila rli it po in t 1 01 2 m pe rle rt ut q aq qa at po in t 1 30 0 m po in t 1 13 7 m po in t 1 53 0 m n ar sa p q aq qa a q in ng us aq po in t 1 57 8 m ik or far su it j e rn py nt enqa sig iss at ki ng itt uu sa q ve st er da len po in t 1 13 2 m sermersuaq ser mers uaq po in t 1 64 0 m ku ug an ng ua q se kv an da le n u iff aqal an ng up q aq qa i fi g. 6 . m ap w ith p la ce n am es u se d in th e t ex t. bl ue d ot s a re th e l oc at io ns o f t he p ro fil es sh ow n in f ig s 7 –1 4. t he ic e c ap s a nd th e d ra in ag e p at te rn g iv e a n im pr es sio n of th e r ug ge d to po gr ap hy an d th e m an y e le va te d ar ea s o f d iff ic ul t a cc es s t ha t d o no t a pp ea r f ro m th e g eo lo gi ca l m ap s i n fi gs 1 an d 4. a dd iti on al n am es fo r n or th -w es te rn d isk o in f ig . 1 03 . 2020 methods detailed descriptions of the methods for field work, photogrammetry and geochemistry are included in pedersen et al. (2017) and are not repeated here. however, the following information is necessary for the independent use of this bulletin. geochemistry: for the maligât formation, 1440 samples were analysed for major elements and 400 for trace elements by x-ray fluorescence spectrometry (xrf); 150 samples were analysed for trace elements by inductively coupled plasma mass spectrometry (icp-ms). for sills and dykes, 560 samples were analysed for major elements by xrf, 170 for trace elements by xrf, and 140 for trace elements by icp-ms. to represent the range of possible contaminants for the magmas, 79 samples of mudstones and sandstones (partly as xenoliths in the igneous rocks) and precambrian basement rocks were analysed for major elements by xrf, 33 for trace elements by xrf and 16 for trace elements by icp-ms. stratigraphic subdivision and numerical coding system the volcanic succession is subdivided into many lithological units both large and small. in order to handle both the units and the large number of samples of the various units, a four-digit sample coding system was established, shown in table 1. the first digit signifies the formation (e.g. 5 for the maligât formation), and the second and third digits signify the unit (e.g. unit 20 in the maligât formation is the nordfjord member); thus all units in the succession have unique three-digit codes. the fourth digit signifies the lithological character of a sample (e.g. 4 for a xenolith). thus, a sample coded 5204 is a xenolith from the nordfjord member of the maligât formation. in the descriptions and figures in this bulletin, individual units are frequently denoted by their identifying threedigit codes, e.g. unit 520 is the nordfjord member. other units are also included in the numerical coding system; intrusion codes have first digit 6 (table 1). the codes covering the vaigat formation have first digit 4 (pedersen et al. 2017, table 1). the numerical coding system was not used in the published geological maps and sections because it was completed after the publication of most of these. the unit codes have been added on many of the excerpts of maps and sections shown in this bulletin. nomenclature the rules of the international stratigraphic guide (murphy & salvador 1999) state that the spelling of a name of an established stratigraphic unit must not be changed even if the spelling of the place name has changed. therefore, the old spellings of the maligât formation and the rinks dal member are maintained although the place names are now spelled maligaat and rink dal. the igneous rocks are named in accordance with the conventional chemical classification of le maitre (2002), based on the total-alkali – silica (tas) diagram. picrites have mgo ≥12 wt%. we use ‘magnesian basalts’ for basalts with 10–12 wt% mgo which are usually visibly olivine-phyric. we use ‘silicic basalts’ for basalts which are crustally contaminated but have sio2 <52 wt% (the boundary to basaltic andesites), because uncontaminated basalts in the nuussuaq basin rarely contain more than 50 wt% sio2. it should be noted that the silicic rocks in west greenland are quite distinct from orogenic rocks with the same tas names. the basaltic andesites and andesites in west greenland have much higher contents of mgo, ni, and cr than their orogenic namesakes, and ‘magnesian basaltic andesite’ and ‘magnesian andesite’ are appropriate but cumbersome names. the lithological terms ‘hyaloclastite’ and ‘pillow breccia’ are frequently used synonymously. following white & houghton (2006), we use hyaloclastite as a general term to denote primary clastic subaqueous volcanic deposits. however, we sometimes use pillow breccia for coarse, clast-rich deposits even though these may include finer-grained, clast-poor parts. the term ‘volcaniclastic’ is used as a nongenetic term to denote any clastic deposit with a large component of volcanic material and includes both primary volcanic and nonvolcanic deposits such as mass-flows and other redeposited sediments. geological sections and profiles in this bulletin, we frequently refer to the five published, photogrammetrically interpreted, long geological sections across the nuussuaq basin (pedersen et al. 2006b). to facilitate reading, they will usually be referred to in short form as follows: “north nuussuaq section” is pedersen et al. (2006a). “central nuussuaq section” is pedersen et al. (2002a). “south nuussuaq section” is pedersen et al. (1993). 21 “central disko section” is pedersen et al. (2005). ”south disko section” is pedersen et al. (2003). all these sections possess a horizontal distance scale along the base (an ‘x axis’) with origin at the westernmost (left) end of the 82–131 km long sections. the scales are preserved in the excerpts shown here, and an expression such as “at 9–10 km” refers to the distance scale. ‘profiles’ are visited, measured and sampled local sections, usually on mountain sides, and are presented in diagrams as vertical columns. the profiles constitute the basis for the definitions and correlations of the volcanic succession. the lithostratigraphic type and reference sections are illustrated by their corresponding profiles. 9 other 4 xenolith 3 plug or feeder dyke 2 sill 1 dyke intrusives: 9 other, incl. zeolites 8 invasive lava or sill 7 pebbles 6 tuff 5 sediment 4 xenolith 3 hyaloclastite matrix 2 pillow 1 lava flow maligât fm: code 5201: maligât fm, nordfjord mb, lava flow example: digit 1: major division digits 2–3: member or unit digit 4: sample lithology 6 intrusives 29 dykes in giesecke dal and inner hammer dal 28 contaminated dykes, unspecified 27 nordfjord complex 26 hammer dal complex 18 weakly contaminated dykes, eastern disko and nuussuaq 17 killiit dyke 16 hanekammen complex 5 maligât fm 99 other 40 sapernuvik member 32 upper niaqussat member 31 middle niaqussat member 30 lower niaqussat member 20 nordfjord member 18 upper rinks dal member, top flows 17 upper rinks dal member, upper main part 16 upper rinks dal member, low-ti flows 15 upper rinks dal member, lower main part 14 upper rinks dal member, upper transition flows 13 middle rinks dal member, akuarut unit 12 lower rinks dal member, lower transition flows 11 lower rinks dal member, skarvefjeld unit 09 lower rinks dal member between low-ti flows and skarvefjeld unit 07 lower rinks dal member, low-ti flows 06 lower rinks dal member between lowest flows and low-ti flows 05 lower rinks dal member, lowest flows 01 rinks dal member, unspecified for definition of the divisions based on tio2 contents, see table 2 and fig. 21. table 1. numerical codes for units and samples of the maligât formation and equivalent intrusions 2222 west greenland basalt group the tertiary volcanic rocks in west greenland are all included in the west greenland basalt group defined by hald & pedersen (1975). the present work does not involve any revision of the original definition of the group as such. the volcanic succession on disko and nuussuaq (with hareøen) east of the itilli fault (figs 1, 2, 4) comprises the vaigat formation and the revised maligât formation and is all of paleocene age. summary of the petrology of the vaigat and maligât formations the petrology of the vaigat and maligât formations was presented by larsen & pedersen (2009), and their conclusions are summarised below. the primary magmas for both formations were highly magnesian, picritic, with at least 16.6 wt% mgo. melting in the asthenosphere took place in garnet facies because the melting column was curtailed by the 100 km thick lithosphere. the vaigat formation magmas fractionated olivine in the feeder channels before eruption but generally did not stop in magma chambers. only minor magma batches stopped in high-level magma chambers where many of them became crustally contaminated. at the time of transition to the maligât formation large, longlived magma chambers developed, presumably in the lower crust. the maligât formation basalts have compositions buffered by magma chamber processes; large-scale cyclicity can be seen, with more and less fractionated steady-state compositions (see fig. 21). repeated pulses of new picritic magmas can be identified in both formations. the sr, nd and pb isotope data show that the asthenospheric mantle source was heterogeneous and the dominant component was depleted, morb-like mantle with nb/la <1. an additional, less-depleted mantle component akin to the iceland mantle source is evident in the upper vaigat formation (ordlingassoq member) but disappeared again at the transition to the maligât formation. an incompatible-element-rich lithospheric mantle component is seen in the alkaline manîtdlat member of the vaigat formation. single scattered tholeiitic lavas somewhat enriched in the same trace elements occur in both formations and are considered to be slightly contaminated with this component. the volume of these rocks is diminutive. in the vaigat formation, crustal contamination was episodic within the short-lived, high-level magma chambers in the sedimentary succession. in the maligât formation, the earliest part, the rinks dal member, is uncontaminated. in the subsequent nordfjord and niaqussat members, weak contamination was continuous in the deep, long-lived magma chambers, possibly because these had moved upwards in the crust, whereas additional episodes of strong contamination took place in local highlevel, subvolcanic magma chambers. the predominant contaminants were the sediments in the nuussuaq basin. when carbon-rich mudstones were assimilated, severe reduction led to formation of graphite or native iron. all rocks more evolved than basalt arose by contamination processes, which will be discussed in more detail in the final chapter of this bulletin. 23 maligât formation revision of the maligât formation the maligât formation was defined by hald & pedersen (1975) with distribution not specified but mentioned on disko, hareøen, nuussuaq, ubekendt ejland, svartenhuk halvø and north of this. the type section was defined along the north and west coasts of nuussuaq ”because of the greater thickness of the formation there as compared with disko”. for the svartenhuk halvø area, the lava succession thought to be laterally correlative with the maligât formation was later termed the svartenhuk formation by j.g. larsen & grocott (1991) and j.g. larsen & pulvertaft (2000). subdivisions of the maligât formation have been made for separate areas: for disko by pedersen (1975a), for western nuussuaq and hareøen by hald (1977) and for ubekendt ejland by j.g. larsen (1977). the svartenhuk formation was subdivided by j.g. larsen & grocott (1991). radiometric age dating of the volcanic succession in the nuussuaq basin by the 40ar−39ar method has revealed the presence of three age groups in the maligât formation as originally defined (fig. 2). throughout disko and nuussuaq east of the itilli fault, the maligât formation is of paleocene age, 61.2−60.3 ma (storey et al. 1998, ages recalculated to an age of 28.201 ma for the fct3 standard, following the geologic time scale 2012). on ubekendt ejland and svartenhuk halvø, the basalts referred to the svartenhuk formation are of latest paleocene age, 60.3−58.0 ma, i.e. younger than the maligât formation on disko and nuussuaq (larsen et al. 2016). on nuussuaq west of the itilli fault, hareøen, western ubekendt ejland and western svartenhuk halvø, an eocene basalt succession is present, dated to 56.2−54.0 ma and recently referred to the naqerloq formation (larsen et al. 2016). this includes the kanísut member of the original maligât formation, which is present on western nuussuaq and on hareøen (hald 1976, 1977) where it overlies sediments with a large volcaniclastic component, suggestive of a break in the volcanic activity. in view of the age differences, we here revise the maligât formation to exclude the eocene part of the succession on hareøen and nuussuaq. as the old type section on western nuussuaq was centred on the eocene lavas, a new composite type section has been defined on central and north-west disko. lithostratigraphy of the maligât formation revised formation history. originally defined by hald & pedersen (1975) to cover disko, hareøen, nuussuaq, ubekendt ejland and svartenhuk halvø. the present revision restricts its known extent to disko and nuussuaq east of the itilli fault. name. after maligaat (old spelling maligât), the strait between disko and hareøen (fig. 4). distribution. the maligât formation extends over the whole of disko and parts of southern and eastern nuussuaq (fig. 4). the depocentre was situated on southcentral disko from where the successive volcanic units spread gradually towards west, east and north. the rinks dal member did not extend north of nuussuaq, and this is most probably also the case for the younger members. the eastern and northern limits are relatively well defined, whereas the southern and western limits are unknown. the formation is assumed to be present on the shelf west and south-west of disko. type section. the type section is composite. for the rinks dal member: the south side of the mountain orpiit qaqqaat at 0–1080 m, central disko (fig. 8, profile 4; fig. 24); the uppermost flows in the member are missing. for the nordfjord member and the lower and middle niaqussat member: point 440 m, northern gully, northwest disko (fig. 14, profile 11; also fig. 104). for the upper niaqussat and sapernuvik members: sapernuvik on central west disko (fig. 14, profile 1; also fig. 161). reference section. the uppermost flows of the rinks dal member and the boundary to the nordfjord member are covered by a short profile on the eastern side of the mountain akuliarusersuaq at 990–1045 m, 13 km south of the type section (fig. 8, profile 3). thickness. the thickness varies from around 2000 m on western disko where the top of the formation is least eroded and the base is below exposure level, through present (eroded) thicknesses of 500–1000 m over large parts of disko and southern nuussuaq, to 200–300 m in the 2424 easternmost parts of disko and nuussuaq where the formation peters out towards the east and north. lithology. the main lithology of the maligât formation is plagioclase-phyric to aphyric tholeiitic basalts that typically form 10–40 m thick lava flows of considerable lateral extent. the major part of the rinks dal member is made up of such flows, which form subaerial lava flows and also subaqueous and invasive flows and hyaloclastites. the rocks of the overlying nordfjord and niaqussat members are more variable and comprise slightly contaminated basalts and picrites and strongly contaminated basaltic andesites, andesites, dacites and rhyolites. these rocks form subaerial lava flows, high-level intrusions and craters, rhyolitic tuff layers and tuffaceous sediments and conglomerates. some of the more contaminated rocks are highly reduced and carry native iron. subdivisions. the maligât formation is subdivided into four members based on lithologies and chemical compositions. these are the rinks dal member of normal tholeiitic basalts, the nordfjord and niaqussat members of weakly to strongly crustally contaminated rocks, and the small sapernuvik member of uncontaminated tholeiitic basalts. the thick rinks dal member is further informally subdivided into 12 chemostratigraphic units. boundaries. the lower boundary is below exposure level on western disko. where exposed, the lower boundary is uncomformable and the maligât formation rests on the disko gneiss ridge on central and south disko, on the lavas of the vaigat formation on northern disko and south and central nuussuaq, on contemporaneous fluvial and lacustrine sediments on eastern disko and southeastern nuussuaq, and on the gneiss highland on eastern nuussuaq. the upper boundary is erosional. the youngest known flows are preserved at sapernuvik on western disko and on the high ground above 2000 m altitude on eastern nuussuaq. age. paleocene, 61–60 ma, magnetochron c26r, based on radiometric dating (storey et al. 1998; larsen et al. 2016). correlation. see fig. 2. no correlation across the itilli fault to hareøen and westernmost nuussuaq has been established. in particular, the relationship to the nûluk member on hareøen and westernmost nuussuaq (hald 1976) is uncertain. on ubekendt ejland, the maligât formation may be time-equivalent to the lower part of the 3000–3500 m thick qeqertalik member. on svartenhuk halvø, the maligât formation is possibly time-equivalent to the sediments of the lowest member of the svartenhuk formation (larsen et al. 2016). general features of the maligât formation subdivision on north-western disko, pedersen (1975a) subdivided the maligât formation into the rinks dal, nordfjord and niaqussat members. this subdivision has been used in many subsequent works on the maligât formation, e.g. pedersen (1977a, b); pedersen & larsen (1987); larsen & pedersen (1988, 1989, 1990, 1992); pedersen et al. (1993, 2000, 2001, 2002a, 2003, 2005). in larsen & pedersen (2009) and in the present work we distinguish a fourth member, the small but distinct sapernuvik member at the top of the formation. we also extend the three main members throughout disko and nuussuaq and subdivide them into a number of informal units. at the same time, we revise the boundary between the nordfjord and niaqussat members. the lithological and compositional variation pattern of the maligât formation is different from that of the underlying vaigat formation, and consequently it is not subdivided in a similar way. in particular, the crustally contaminated and uncontaminated rocks do not form interspersed units, as seen in the vaigat formation, but the crustally contaminated rocks follow above the uncontaminated rocks. because of the lithological uniformity of large parts of the succession, in particular the rinks dal member, much of the subdivision is based entirely on geochemical differences. this has required analysis of many flow-by-flow sampled profiles and several samples outside profiles. as the geochemical variations are commonly gradational, the assignment of a flow to a specific unit is subjective in some boundary cases, and most units cannot be mapped in the field. for this reason, most of the units of the three main members of the maligât formation are maintained as informal units. on the other hand, several of the units are of large lateral extent because the major part of the formation is in subaerial lava facies. this is in contrast to the strong lateral variation and progradation of the hyaloclastites and subaerial lava flows of the successive units in the vaigat formation (pedersen et al. 2017). 25 the complete subdivision of the maligât formation used in this bulletin is shown in table 1. this table shows the 4-digit sample coding system used in order to handle the large number of samples of the various lithological units (see chapter on methods above). all geochemical analyses of the rocks of the maligât formation, their xenoliths, associated dykes and sediments representing possible contaminants are available in a supplementary data file at www.geus.dk/bulletin40. distribution of units. the location of the major profiles that form the basis for the mapping of the maligât formation is shown in figs 4 and 6. stratigraphic panels with the subdivided profiles are shown in figs 7–14 which cover 1130 out of 1450 analysed samples. several short profiles, particularly on western disko, are not shown. successive stages in the deposition of the maligât formation are shown in fig. 15. the oldest known parts are found on southern disko on and near the disko gneiss ridge. from here, the successive volcanic units spread northwards, overlapping the low picrite shield of the vaigat formation, and eastwards, filling in the assoq lake (see below). the middle part of the rinks dal member (the akuarut unit) reached southern nuussuaq, and the upper part of the rinks dal member extended to northeastern nuussuaq and transgressed the eastern boundary fault there. in the easternmost lava exposures on northern nuussuaq, lavas of the niaqussat member rest directly on the gneiss or are invasive into sediments of the atanikerluk formation. lithological variations and the filling of the assoq lake in order to provide a framework for understanding the lithological variations described in the following, particularly in the rinks dal member, the history of the assoq lake is summarised here. the existence of a large lake on central and eastern disko and southern nuussuaq is indicated by the wide extent of a succession of non-marine mudstones coarsening upwards into sandstones; these sediments form the assoq and umiussat members of the atanikerluk formation (larsen et al. 2006; dam et al. 2009, fig. 131). the history of filling of the lake by contemporaneous volcanic rocks and sediments is revealed by the distribution of lavas in subaerial, subaqueous, water-influenced and invasive facies and their relations to the sediments. the various volcanic facies and their diagnostic characters are as follows. subaerial lava flows have masive central parts and vesicular, slaggy or ropy top zones that commonly, but not always, are red-oxidised. in places, there are brick-red laterite horizons between flows. flows emplaced over wet surfaces or into shallow water have regularly columnar-jointed zones (colonnades) at the base and commonly zones with small, irregular, blocky columns (entablatures) in the centre. the vesicular top zones may be red-oxidised. flows emplaced into deeper water at high rates have colonnades and entablatures and thick to very thick, glassrich, brecciated top zones; if the water depth was larger than the thickness of the flows the tops are yellowish brown and there is no red-oxidation. flows emplaced into deeper water at low rates are brecciated throughout and form hyaloclastite deposits, which are commonly foreset-bedded; if these are capped by subaerial flows from the same eruption the height of the foresets is a reliable measure of the water depth. examples of these four types of flows are seen in the coastal cliff beneath skarvefjeld (see fig. 56). invasive lava flows are subaerial flows that have met unconsolidated sediment during their course and continued their flowage as sheets intruding laterally into the sediment, in places for distances of several kilometres. such flows have the appearance of sills in the sediment but are stratigraphically in place in the volcanic succession (figs 7–11) and can in some cases be followed laterally into normal subaerial facies. invasive flows are usually waterinfluenced with colonnades and entablatures, in places double colonnades (see e.g. fig. 42); their tops are chilled against the overlying sediment and have commonly developed small lobes intrusive into the sediment (see fig. 76). the overlying sediment may be only a few metres thick, in which case the lava flow may have extruded up through the sediment to produce rootless craters, as at marraat qaqqaat (see fig. 70). the flows may invade either mud or sand horizons; mud horizons seem to be more prone to invasion than sand horizons. invasive lava flows were first described from the columbia river basalts by schmincke (1967) and later by e.g. byerly & swanson (1978) and ross (1989). the assoq lake, or possibly system of lakes, presumably had its largest extent at the onset of extrusion of the maligât formation (fig. 16); it was bounded to the west by the disko gneiss ridge, and to the north it covered the edge of the lava shield of the vaigat formation, whereas its eastern boundary was most probably towards a low plain formed by siliciclastic sediments sourced from the east and south. its southern boundary is unconstrained 2626 fig. 7 (two pages). profiles through the maligât formation on southern and eastern disko. the legend applies to all profiles in the maligât formation in subsequent figures. 518 517 516 515 520 530 514 513 512 511 509 512 512 514 513 515 517 518 509 506 509 512 513 514 515 516 517 509 507 507 50 7 506 326406 326407 326408 326409 saprolite gneiss 326405 340929 340930 340931 340932 340933 340934 340935 340936 340937 340938 340939 340940 340941 340942 340943 340944 340928 340927 340926 340921 340922 340923 340924 340925 340952 dinos dinos 340951 340950 340945-40 tuff layers 340953 340956 328419 328420 328421 328431 327011 327012 327013 327014 327015, 17 327016 327018 327019 327020 157261 157262 157263 157264 157265 157266 157267 157268 157269 157270 157271 157272 157273 157274 157275 157276 157277 157278 157279 157280 327021 327022 327042 327023 327043 327045 327046 327047 327048 327049 327050 327051 327056 327057 327058 327059 327060 327061 327064 327065 327066 157254 157252 157250 157246 157249 157244 157241 157238 157235 157234 327062 327063 sand 328432 328433 328434 328435 328436 328437 328438 328439 328440 328441 328442 328443 fault 328454 328453 328452 328451 328450 328449 328448 328447 328446 328445 328444 fault rootless cones 328474 328473 328472 328471 328470 328469 328468 328467 328466 328465 328464 326419 326418 326417 326411-13 326410 326420 326416 326414 326415 326430 326431 326432 326433 326438 326439 327069 327070 327071 327072 327073 327074 327075 327076 327077 327078 327079 327080 327081 327082 279055 279054 279053 327068 327067 sand shale 700 750 600 650 800 850 900 m 350 400 250 300 450 500 50 100 0 150 200 550 350 300 600 650 m 400 250 300 450 500 200 150 550 alanngup qaqqai 1 uiffaq cliff 5 brededal gv04_06_016_lml southern disko (west) w e 50 100 0 150 350 400 250 300 450 500 m 50 100 0 150 200 4 innarsuaq–ippik skarvefjeld 2 killiit fortunebay 2 killiit fortunebay 700 750 600 650 800 m 500 550 7 marraat qaqqaat 6 assoq (stacked stratigraphy in slipped blocks) 700 750 600 650 800 850 900 m 350 400 250 300 450 500 50 100 0 150 200 550 700 600 650 350 400 250 300 450 500 50 100 0 150 200 550 850850 50 100 0 150 200 3 akuarut lyngmarksfjeld tuapassuit 27 518 520 518 517 520 530 516 518 515 514 512 516 515 514 513 512 511 517 516 515 514 513 512 509 327010 327009 327008 327007 326798 326799 327001-02 327003 318842 slipped 327005 323317 dinos 318841 318840 318839 318838 318837 318836 318835 318834 318833 318832 318831 327006 318868 318867 318866 318865 318864 318863 318845 318844 318848 318849 cgl. 326791 326792 + tuff + tuff 326793 340920 340919 340918 340917 340916 340915 coal coal coal coal slipped megablock 340914 340913 340912 340911 340910 340909 340902 340903 340904 340905 340906 340907 340908 362164 362165 362166 362167 362168 362343 362344 362345 362346 362347 362348 breccia laterite short columns or flow folding top rubble entablature zone colonnade pahoehoe lava flows mudstone and siltstone sandstone not exposed basement gneiss no data conglomeratecgl. native-iron-bearingfe enriched in some trace elementse lithological unit code509 dinoflagellate cystsdinos foreset-bedded hyaloclastites coarsely columnar-jointed lava flow 318808 318809 318810 318811 318812 318813 318814 318815 318816 318817 318818 318819 318820 pillowed top 354799 354798 156759 156760 156757 318790 318789 318788 318787 318786 coal 318785 318784 318791 362348 sample legend for all maligât formation profiles fe 650 m 400 450 300 350 500 550 600 m m a bo ve se a le ve l ri nk s d al m em be r n or df jo rd m em be r n iaq us sa t m em be r sapernuvik mb 1000 m 750 800 850 900 950 950 m 700 750 600 550 650 800 850 900 950 m 700 750 650 800 850 700 750 800 850 900 1000 950 700 750 650 800 850 900 1000 1100 m 1050 8 tuapaat qaqqaat 9 point 1025 m 12 point 1123 m 13 aqajaruata qaqqaa 10 skorstensfjeld 11 sortebærdalen point 1014 m 950 700 750 600 650 800 850 900 1000 1050 m 700 750 600 650 800 850 m 14 inngigissoq gv04_06_015_lml southern disko (east) sw ne 70°n disko 50 km 14 13 12 11 10 9 8 7 4321 5-6 53°w 512 511 509 507 506 505 518 517 516 515 514 513 dacite andesite basaltic andesite basalt alkali basalt upper middle basaltic andesite lower 505 506 507 509 511 513 514 515 516 517 530 531 511 512 512 518 516 517 515 513 514 513 512 513 515 516 517 518 520 514 362352 175041 175042 175043 175044 175046 175073 175074 175075 175077 175078 175079 175080 175081 175082 175048 175049 175052 175054 175055 175056 175058 175060 175061 175063 175067 pillow lava ponded arkose/cgl. 328490 328489 328488 328487 328486 328485 328484 328483 328482 328481 328480 328479 354704 354705 354706 354707 354708 354709 354710 354711 354712 354713 354714 354715 354716 354717 354718 354733 354732 354731 354730 354727 354726 354725 354724 354723 354722 354721 354719-20 354728-29 cgl. 328491 328478 328477 328476 328475 328492 328493 328418 328405 328406 pegmatite 328407 328408 328409 328410 328411 328412 328413 328414 328415 328403 333000 332999 332998 332997 332996 332995 332994 332993 332992 332991 332990 332989 332988 332987 332986 328402 328401 328417 328404 362309 362310 362311 362312 362313 362292 362293 362294 362295 362296 362297 362298 362299 362300 362301 362302 362303 362315 362316 362291 362305-07 sandstone? sandstone? projected 362353 > five flows unsampled 354797 354796 362336 362335 362334 362333 362332 362331 coal 362330 362329 coal 362328 362327 362326 362319 362320 362321 362322 362323 362318 gv04_06_017_lml central disko w e 700 750 600 650 800 850 900 950 1000 1050 1100 m 350 400 300 450 500 550 1000 1050 m 700 750 600 650 800 850 900 950 1000 1050 1100 1150 1200 m 350 400 450 500 550 700 750 600 650 800 850 900 950 1000 1050 350 400 250 300 450 500 100 150 200 550 700 750 600 650 800 850 900 950 1000 1050 1100 350 400 250 300 450 500 100 150 200 550 700 750 650 800 850 900 m 700 750 600 650 800 850 900 950 1000 1050 1100 m 2 eqalunnguaqqat qaqqaat 4 orpiit qaqqaat 3 akuliarusersuaq 1 tunup qaqqaa 5 daugaard-jensen dal 6 laksedalen w 7 blåbærdalen 50 100 0 profile shift profile shift 175096 175099 175102 175103 175104 175105 175108 175109 175113 175114 175115 175116 175118 175120 600 350 400 250 300 450 500 100 50 0 150 200 550 disko 70°n 50 km 6 54 3 21 7 53°w fig. 8. profiles through the maligât formation in a w–e panel across south-central disko. the orpiit qaqqaat profile (no. 4) represents the type section for the rinks dal member. for legend, see fig. 7. 29 517 518 520 530 517 517 515 513 513 511 511 511 511 512 512 512 517 520 530 532 513 511 515 514 513 512 511 509 509 509 507 507 513 513 513 514 512 511 507 506 505 516 506 507 509 340929 340930 340931 340932 340933 340934 340935 340936 340937 340938 340939 340940 340941 340942 340943 340944 340928 340927 340926 340921 340922 340923 340924 340925 340956 328490 328489 328488 328487 328486 328485 328484 328483 328482 328481 328480 328479 354704 354705 354706 354707 354708 354709 354710 354711 354712 354713 354714 354715 354716 354717 354718 354733 354732 354731 354730 354727 354726 354725 354724 354723 354722 354721 (cut off) 354728-29 cgl. 328491 332939 332940 332941 332942 acid tuff 332944 332943 332945 332946 332947 332948 332949 332950 plateau edge 332951 un sa m pl ed in te rv al 332905 tuff 136951 136952 136953 136954 136955 136956 136957 136958 136959 136961 332907 332908 332909 sediment 332910 332911 332912 332913 332914 332915 332916 332917 332918 332919 332920 332921 332922 332923 332924 332926 332927 332928 332929 332930 332931 332932 332861 332862 332863 332865 332866 332867 332868 332869 332871 332872 332873 332874 332875 332876 332877 332878 135996 135997 135998 135999 136900 136901 136902 e 136967 e 136968 e 136969 e 136970 e 136971 e 136972 136973 136974 vaigat formation vaigat formation 136903 profile shift 332879 332880 332881 332925 332952 332984 332982 332981 332980 332979 332978 332977 332976 332975 332974 332973 332972 332933 332962 332961 332960 332959 332958 332957 332956 lava tongues 332955 332954 332937 332934 332935 mudstone 332983, 953 328478 328477 328476 328475 328492 328493 700 750 600 650 800 850 900 m 350 400 250 300 450 500 50 100 0 150 200 550 1 brededal gv04_06_019_lml central disko s n 850 1000 1050 m 700 750 600 650 800 850 900 950 1000 1050 350 400 250 300 450 500 100 150 200 550 850 3 orpiit qaqqaat 2 akuliarusersuaq 4 north of sorte hak 5 kuugannguaq se 6 pyramiden 7 orlingasoq 700 750 600 650 800 850 900 950 1000 1100 1200 1300 1050 1150 1250 1350 m 350 400 300 450 500 550 1400 1500 1600 1450 1550 1650 750 800 850 900 950 1000 1100 1200 1300 1050 1150 1250 1350 1700 1800 1900 m 1750 1850 1500 1600 1550 1650 1500 m 1200 1300 1400 1250 1350 1450 850 250 300 200 70°n disko 50 km 1 2 3 4 5 6 7 53°w fig. 9. profiles through the maligât formation in a s–n panel across disko. for legend, see fig. 7. 3030 513 513 517 513 520 530 530 509 511 511, 512 511 511 511 512 509 512 517 517 516 515 514 513 512 511 509 279036 279033 279032 279031 279030 279029 401516 401515 401514 401513 401512 401511 401510 401509 136944 e 136945 e 136946 136947 136948 e 136949 136950 vaigat formation 136951 136952 136953 136954 136955 136956 136957 136958 136959 136961 279037 362377 279097 279098 279099 invasive 279100 rhyolite tuff cgl. unsampled interval rhyolite tuff cgl. 279401 279402 279403 362175 362172 362171 362174 362173 362376 362375 362374 362373 362372 362371 362370 362369 362368 362367 362360 362361 362362 362366 279038 279039 279040 279041 279042 279043 279044 279045 279046 279047 279048 279049 279050 279051 279052 401507 401506 401504 vaigat formation profile shift 156730 tuff tuff 135996 135997 135998 135999 136900 136901 136902 e 136967 e 136968 e 136969 e 136970 e 136971 e 136972 136973 136974 vaigat formation 136903 profile shift gv04_06_020_lml north-eastern disko (north-west)nw se 700 750 650 800 700 750 600 650 800 850 1000 950 900 1050 1100 1150 1200 1300 1400 1500 1250 1350 1450 1550 850 900 950 1000 1100 1200 1050 1150 1250 m 2 asuutaa 5 inussuk 4 killerpaat qaqqarsuat qullissat 7 narsap qaqqaa 3 pyramiden 6 point 1530 m 1700 1800 1900 m 1750 1850 1500 1600 1550 1650 1150 1200 1050 1100 1300 1 orlingasoq 1500 m 1200 1300 1400 1250 1350 1450 70°n disko 50 km 1 3 2 4 5 6 7 8 9 10 12 11 53°w fig. 10 (two pages). profiles through the maligât formation on north-eastern disko. for legend, see fig. 7. 31 518 520 530 531 532 517 516 515 514 513 514 515 516 517 518 520 512 513 512 511 340816 340817 340818 340819 340820 340821 340822 340823 340826 340827 profile shift 340828 318826 318827 318828 318829 318830 354765 354766 354767 354768 354769 354770 354771 354773 354774 354775 354776 362357-58 fe 340829 340830 340831 340832 340833 340834 340835 340836 340837 340863 340864 340865 340866 340867 340849 340850 sediment sediment340851 340852 340824-25 362176 362178 362177 362179 fe 362399 362400 sediment 362181, 349901 362180, 349902 362398 362393 362394-97 rhyolite tuff 362392 362391 362389 sediment 362386 sediment 362388 362385 362384 340900 340899 340898 340897 340896 340895 sediment 340894 340893 340892 340891 340890 340889 340888 340887 340873 340882 340883 340885 340886 340870, 74, 75 hyaloclastite 340877, 84 340901, 362383 340853 340857 340858-59 340860-61 340862 340868 340845 340846 340847 340848 340844 340843 340840 340839 340841-42 340854-56 dinos 354777 354778 354779 slipped gv04_06_024_lml north-eastern disko (south-east)wnw ese9 frederik lange dal charles polaris dal 8 qinngusaq inner kvandalen 10 kvandalen 700 750 600 650 800 850 900 950 1000 1100 1200 1300 m 1050 1150 1250 450 500 550 850 700 750 600 650 800 850 900 950 1000 1100 1200 m 1050 1150 350 400 450 500 550 850 700 750 600 650 800 850 900 950 1000 1100 1200 1300 m 1050 1140 1150 1250 450 500 550 850 + coal 318808 318809 fe 318810 318811 318812 318813 318814 318815 318816 318817 318818 318819 318820 pillowed top assoq mb 354799 354798 156759 156760 156757 318790 318789 318788 318787 318786 318785 318784 318791 11 aqajaruata qaqqaa 950 700 750 600 650 800 850 900 1000 1050 m 700 750 600 650 800 850 m 12 inngigissoq 3232 514 517 520 516 513 517 516 515 514 513 514 513 514 513 514 516 511 513 514 515 516 517 515 514 vaigat formation vaigat formation 400265 400264 362212 362213 362214 vaigat formation 362211 362210 m? 362209 332843 332845 332846 332847 332848 332849 332850 m 332851 332852 362195 362196 m 362197 e 402550 402551 402538 402539 402540 402541 402542 402543 402544 402545 402546 402547 402548 402549 vaigat formation sediment sediment sediment 362198 e 362199 362201 362235 assoq mb vaigat formation 362194 362193 332853 e 332854 e 332855 e vaigat formation 362048 400263 400262 m m : thick marker flow e : enriched chemistry 400261 400260 400259 400258 400257 400256 400255 400254 400253 400252 400251 400250 400249 sediment gv04_06_021_lml s. nuussuaq (north-west)nw se 1750 m 1700 1550 1600 1650 1700 1750 1800 1850 1900 2 point 1888 m 1400 1450 1500 1550 1600 1650 1700 1750 1800 4 ivissussat qaqqaat 5 ridge s of point 1760 m 6 eqi 3 point 1722 m 1 niiortuut 1500 1550 1600 1650 1200 1250 1300 1350 1400 1450 1500 1550 1600 1200 1250 1300 1350 1400 1450 m fig. 11 (two pages). profiles through the maligât formation along the south coast of nuussuaq. for legend, see fig. 7. 33 517 516 515 514 513 514 513 512 511 514 514 513 512 515 516 517 517 520 318749 318750 318751 sediment 318759 sediment 318754 318763 gneiss 318764 318765318770 tuff?318768 318766 318769 318755 318756 362227 362226 e 318758 e 362225 362222 362220 362219 +sandstone 362218 489127 naujât mb mudstone 489108 489103 e 489104 traces of sediments 362185 362184 362183 362182 362207 362191 362192 318703 assoq mb assoq mb 318702 318704 318705 362208 362186 362187 362188 362189 362190 489105 489106 489107 318748 318747 318746 318745 318762 no samples 318761 318760 sediment sediment sediment m gv04_06_022_lml s. nuussuaq (south-east)nw se 7 giesecke monument uppalluk 8 eqip qaqqaa 9 umiusat 10 tartunaq 11 saqqaq 1200 1250 1100 1150 1300 1350 1400 1450 1500 1550 1600 m 850 900 950 1000 1050 1200 1100 1150 850 800 750 900 950 1000 1050 1200 1250 m 1100 1150 m 850 800 750 900 950 1000 1050 850 800 750 900 950 disko 70°n 53°w 50 km nuussuaq 1 2 3 4 5 6 14 12 13 8 7 9 10 11 3434 (dam et al. 2009). scarce marine dinoflagellate cysts in samples from eastern and south-eastern disko indicate that the lake was subject to marine inundations, most probably from the south (dam et al. 2009). the western shoreline towards the disko gneiss ridge and later volcanic rocks was probably steep; c. 200 m of foresetbedded hyaloclastites at orpiit qaqqaat just east of the gneiss ridge (fig. 8, profile 4) suggest early water depths of this magnitude. the water depths shallowed towards east and north, and the north-western shoreline towards the vaigat formation and the eastern and northern siliciclastic shorelines all had low gradients. the height of the hyaloclastite foresets of the skarvefjeld unit (511) indicates the depth of the water at one point in time during the filling of the lake: up to 100 m at skarvefjeld, 60 m in blåbærdalen, 60 m in frederik lange dal, at least 90 m at qinngusaq (inner kvandalen), more than 30 m at inussuk on north-east disko, and 0 m at qullissat only 4–5 km farther north where the corresponding flows are subaerial. extensive peat swamps developed along the eastern shoreline (eastern disko). as shown in figs 15 and 16 the filling of the assoq lake began on southern disko on and near the disko gneiss ridge. subaqueous lava flows and hyaloclastites of the lower rinks dal member (units 505–511) prograded eastwards and northwards into the lake (figs 7–10). the voluminous unit 509 pushed the western shoreline 10–20 km eastwards, and the hyaloclastites of unit 511 pushed it farther 10–40 km eastwards (fig. 16). the filling tended to decrease the accomodation space and the water level may have risen, but as the basin floor also subsided the balance was delicate. in all, the lake gradually became shallower. at the eastern shore, the sediment accumulation kept pace with the aggrading lava plateau, or outpaced it. around the transition to the middle rinks dal member (unit 513, the akuarut unit), there was apparently a decrease in the magma production rate because sediments, particularly sands, spread widely across the lava plain as far west as to skarvefjeld (figs 7, 8, 10). unit 512 and the thick unit 513 are in subaerial lava facies over large parts of disko, but on eastern disko where the number of flows is strongly reduced, the flows commonly invaded the rising sediment pile as invasive flows. units 512 and 513 were also the first flows of the maligât formation to reach southern nuussuaq where they flowed northwards and occur, successively northwards, in invasive, near-shore and subaerial facies (fig. 11). after the deposition of unit 513, the assoq lake was reduced to small temporary pools, coal swamps and fluvial plains. the flows of the upper rinks dal member (units 514–518) are everywhere subaerial, but on easternmost 517 520 532 532 530 400346 gneiss 400347 400348 400349332893 332894 332895 sediment yellow sediment yellow sediment 332860 332859 332882 332884 tuff tuff 332885 332886 332888 332889 332890 332891 332892 gneiss 332897 332898 332899 332900 gv04_06_023_lml central nuussuaq w e 12 point 2080 m 14 point 1640 m 13 nunavik 1850 1900 1950 m 1500 1550 1400 1450 1600 1700 m 1650 1350 1300 1350 1250 1400 1450 1500 1550 1600 fig. 12. profiles through the maligât formation on central eastern nuussuaq. profile locations in fig. 11. for legend, see fig. 7. 35 517517 515 513 513 512 512 511 517 509 512 511 514 515 518 520 516 515 514 513 512 512 175011 326823 326827 326828 326829 326830 326831 326832 326833 326834 326835 326836 326837 326838 326839 326840 326841 326842 326843 326844 326845 326846 un sa m pl ed in te rv al 264461 264459 264458 264457 264456 264455 264454 264452 175198 175195 175194 175193 264462 326847 326824 326825 326826 264301 156588 156589 156590 156591 more flows more flows more flows more flows several thin, grey flows 156587 156593 156594 326925 high ti 326924 326923 326922 326921 326920 326919 326918 326917326916 326915 326914 326913 326912 326911 326910 326909 326908 326902 326903 326904 326905 326906 326907 fault vaigat fm picrite lavas vaigat fm picrite lavas 156595 156596 156597 156598 156599 156600 156601 156592 264303 264305 264308 264310 264311 264312 175012 175013 175014 175015 175016 175017 175018 175019 175020 175021 175024 175026 175027 175028 175029 175030 350 400 250 300 450 m 50 100 0 150 200 700 m 600 650 350 400 250 300 450 500 100 150 200 550 700 600 650 350 400 250 300 450 500 100 50 150 200 550 700 600 650 550 350 400 250 300 450 500 100 0 50 150 200 550 m 1450 1550 m 1350 1400 1500 1100 1150 1000 1050 1200 1250 850 900 950 1300 1450 m 1350 1400 1100 1150 1000 1050 1200 1250 850 750 800 900 950 1300 1 saqqarliit ilorliit mellemfjord 6 mouth of giesecke dal (e) 5 easternmost hammer dal point 1510 m 4 inner stordal point 1578 m 2 perlertut qaqqaat nordfjord 3 inner nordfjord point 1132 m gv04_06_026_lml s nwestern disko disko 70°n 50 km 2 3 4 1 5 6 53°w fig. 13. profiles through the maligât formation (mainly the rinks dal member) in a s–n panel on western disko. for legend, see fig. 7. 3636 518 520 530 532 520 530530? 531 532 540 532 176613 dyke 263976 263975 263974 263973 263972 263971 263920 264078 264057 264058 264059 264060 264062 264063 264064 fe 264072 264065-67 fe 264070-71 264068-69 fe 264056 264079 264080 264081 264082 264084 264085 264086 264088 fe 264087 263919 263918 263917 263916 263915 263914 263913 263922 263921 263906 263904 263907-08 fe fe 263910-11 263970 263969 263968 263967 263966 263965 263954 profile shift faults profile shift 263956 263957 263958 fe 263961 263960 176611 176610 dyke 176609 176590 tuff tuff tuff tuff 176591 176592 176593 176594 176595 176596 176597 176598 176599 176600 176572-75 fe 176571 tuff 176608 176614 176615 176616 176617 176618 176619 176620 176621 176622 176623 176604 176607 dyke 176602-03 176605 176606 176624 tuff tuff tuff tuff gv04_06_028_lml 550 600 m 300 350 200 250 400 450 100 50 150 500 650 550 600 300 350 200 250 400 450 50 0 100 150 500 1050 m 950 1000 800 850 900 650 m 550 600 400 510/415 550 300 350 410/590 450 250 200 150 500 200 240 100 150 1100 1150 1000 1050 1200 1250 850 900 950 1300 m 50 50 0 100 120 0 s nwestern disko (south) 1 sapernuvik 3 ikorfarsuit mellemfjord 6 qasigissat sælbugten 4 vesterdalen point 1012 m 5 vesterdalen point 1300 m 2 eqaluit nordre laksebugt fig. 14 (two pages). profiles through the maligât formation (mainly the nordfjord and niaqussat members) on westernmost disko. for legend, see fig. 7. all profiles are corrected for dips and faults to show the real thicknesses. for profiles 1–5 the heights are approximate m a.s.l.; the heavily faulted profiles 6–12 are reconstructed to local baselines and correlated relative to the base of the niaqussat member. 37 530 520 518 517 518 520 530 531 113454 crater breccia and lava flow fed from crater 326543 326542 326541 326540 326539 326538 326537 326513 326481 326482 326531 326532 326533 326534 326512 326511 326510 326509 326508 326483 326472-75 326536, 50 cgl. 113459 fe 326643, tuff 326658 326708 326709 326659 cgl. 326660 326663 326704, 5, 6 326703, tuff tuff cgl. 326647, 50 326653, 55 113456 113458, 176410 113463, 176411 crater floor unexposed 274422 274421 274415 274414 274416 274417 274418-20 274413 274412 274411 sediment 274410 274409 274408 274428 274430 274432 274433 176640 176641 176642 176639,38 176643 176637 176636 cgl., tuff 176635 tuffs tuff tuff fe tuff 279475 176443 326714 176492 176493 176497 176491 176490 lava flows filling crater depression massive feeder body in crater breccia crater breccia fe 326715 fault faults fault 326716 326717,18 326719 176447 176448,49 fe 176450 176442 blocky lava top fe fe fe fe 176439 176440,41 176444-46 279476 279484-86 cgl. 279477 279478 279479 279480 279481 279482 279483 279465 279463 279462 156668 279461 156669 156670 156671 156672 156673 156674 279464, 156667 fe 176633 176631 176629 176627 264174 264175 264176 264177 264178 264179 264180 264181 176634 sediment tuff tuffs tuff sediment sediment crater scoria fe fe fe fe tuff tuff sediment sediment, tuff 450 500 550 m 350 400 300 7 kingittuusaq outer nordfjord (south) 8 point 1070 m outer nordfjord (north) 9 point 600 m north side of hammer dal 10 point 440 crater north of hammer dal southern gully 11 point 440 m north of hammer dal northern gully 12 niaqussat gv04_06_027_lml 450 500 m 200 250 100 150 300 350 0 50 400 200 m 100 150 0 50 200 250 100 150 0 50 200 m 100 150 0 50 100 150 m 0 0 50 100 150 m 0 50 s nwestern disko (north) 70°n disko 50 km 8 76 3 4 5 1 2 9 12 10-11 53°w ? ? ? ? ? ? ? ? ? ? no no no no no nq ? ? ? ? vaigat malig ât hareøen 69°30’n 70°30’n 70°n 55°w 53°w 52°w nuussuaq disko vaigat malig ât hareøen subaerial invasive su ba qu eo us su baq ue ou s su bae ria l 69°30’n 70°30’n 70°n 55°w 53°w 52°w nuussuaq vaigat malig ât hareøen 69°30’n 70°30’n 70°n 55°w 53°w 52°w nuussuaq disko vaigat malig ât hareøen 69°30’n 70°30’n 70°n 55°w 53°w 52°w nuussuaq disko vaigat precambrian basement 25 km malig ât hareøen 69°30’n 70°30’n 70°n 55°w 53°w 52°w vaigat formation paleocene sedimentsunknown nuussuaq 507 509505 + 506 505 506 513 514511 518 520+530–532517 disko vaigat malig ât hareøen 69°30’n 70°30’n 70°n 55°w 53°w 52°w nuussuaq disko gv01_02_137_lml vaigat malig ât hareøen 69°30’n 70°30’n 70°n 55°w 53°w 52°w nuussuaq disko vaigat malig ât hareøen 69°30’n 70°30’n 70°n 55°w 53°w 52°w nuussuaq disko vaigat malig ât hareøen 69°30’n 70°30’n 70°n 55°w 53°w 52°w nuussuaq disko 39 disko and north-eastern nuussuaq they are still waterinfluenced and in some places ponded. remnants of the lake may have existed east of the present exposures. later lava flows of the nordfjord and niaqussat members are subaerial except for the uppermost flows of the niaqussat member in the easternmost profile on northeastern nuussuaq. these flows are intercalated with nonmarine sediments and are water-influenced and partly invasive (fig. 12, profile 14; see also figs 164, 165). the sediments are referred to the atanikerluk formation and may have been deposited in the youngest remnant of the assoq lake or in an independently formed local lake in a depression on the gneiss highland. boundaries lower boundary the lower boundary of the maligât formation is everywhere unconformable. the formation overlies precambrian gneiss on southern and central disko and eastern nuussuaq, volcanic rocks of the vaigat formation on north-western and northern disko and central nuussuaq, and sediments of the atanikerluk formation on central and eastern disko and south-eastern nuussuaq (fig. 15). in many places, the boundary is not exposed because it is either covered, below sea level or above erosion level. on south-western disko the character of the substrate is unknown because it is everywhere below sea level. precambrian gneiss. the disko gneiss ridge on southern and central disko has an irregular surface with ‘hilly relief ’ created in the cretaceous by deep weathering and subsequent stripping of material, as described in detail by bonow (2005). the local altitude differences between palaeo-hills and -valleys are to up to 400 m, and the longexposed gneiss surface is in many places covered by thick residual deposits (saprolite). the successive volcanic units of the maligât formation gradually drowned this irregular surface, and at the time of deposition of the middle part of the rinks dal member, all the gneiss hills were covered (fig. 15). the relations between gneiss and volcanic rocks are well illustrated in the south disko photogrammetric section. a few examples are described below. at the entrance to the narrow canyon that separates navaranaat/apostelfjeld and akuarut/lyngmarksfjeld, a weathered and eroded gneiss surface is directly covered by a basaltic lava flow of unit 507 in transitional subaqueous–subaerial facies (fig. 17), see also fig. 36 and bonow (2005, fig. 15). facing page: fig. 15. cartoons showing the substrate and extent of the successive units of the maligât formation. the lithological unit codes are indicated in the lower left corner of each cartoon. units 512 and 515–516 are not shown; their extents are similar to those of units 513 and 514, respectively. substrate in pale colours, legend in cartoons for units 505–507; the largest hills on the disko gneiss ridge are shown with contour lines. the presence of a unit in a profile (figs 4, 6) is marked with a black dot; the observed absence of a unit is marked with an ×. the dotted lines for units 507, 509 and 511 show the position of the western shoreline of the assoq lake at the beginning of the infill of each unit into the lake. units 505 and 506 are mainly in subaqueous facies except on the gneiss ridge, and units 513 and younger are mainly in subaerial facies, changing to invasive in the easternmost areas. ? ? ? ? ? ? 69°30’n 25 km 70°30’n 70°n 55°w 53°w 52°w 50 9 50 7 51 1 51 3 fig. 16. the extent of the assoq lake (blue) during its early stage when it probably reached its largest extent, and during stages of deposition of the maligât formation. substrate colours as in fig. 15. the infill of the four units 507, 509, 511 and 513 of the maligât formation pushed the western shore of the lake successively eastwards: the four dashed lines show the position of the western shoreline at the end of the infill of each unit into the lake. compare with the cartoons in fig, 15 where the shorelines at the beginning of the infill of the units are shown. black dots indicate the sample profiles in figs 7–11 used to delimit the lake. 4040 the contact is well exposed near tuapassuit 3−4 km farther west (south disko section at 74.1−74.2 km, see also fig. 36). the deeply weathered gneiss surface is covered by 3−4 m of saprolitic sediment and overlain by subaerial basaltic lava flows from the maligât formation unit 506 (fig. 7, profile 3; figs 18, 19). in kuannersuit sulluat (the easternmost part of kangerluk/diskofjord) the contact is well exposed on the west coast below eqalunnguaqqat qaqqaat in one of the gullies below point 1109 m (fig. 8, profile 2). a strongly weathered gneiss is gradually transformed into metre-sized blocks packed in a matrix of kaolinised arkose (f. ulff-møller, unpublished field notes 1978). the upper part of the sediment contains layers of mica-rich sandstone and arkose with inter-layered reddish, clay-rich areas. the sediment is about 10 m thick and is covered by a basaltic pillow lava of unit 505. on nuussuaq the maligât formation overran the high gneiss terrain east of the saqqaq–ikorfat part of the eastern boundary fault. north of saqqaq a few small remnants of lava flows (of unit 517) rest on the gneiss and represent an originally much more widespread lava cover (pedersen et al. 2007a; fig. 11, profile 11). there are no well-preserved contacts. north of the aaffarsuaq valley and east of the saqqaq– ikorfat boundary fault the maligât formation covered first the vaigat formation and farther east came into direct contact with the weathered gneiss, which had a considerable topography (central nuussuaq section at 64.5 to 71 km; pedersen et al. 2007a; fig. 12, profiles 13, 14). only a few of these localities have been visited in the field. vaigat formation. on northern disko and southern and central nuussuaq, lava flows of the maligât formation overlie lava flows of the ordlingassoq member of the vaigat formation, either directly or separated from them by a few decimetres to metres of residual soil with varying amounts of plant remains. at any one locality, the contact between the two formations is simple and subhorizontal (fig. 20). however, because of the shield shape of the vaigat formation and also subtle syn-volcanic basin movements, the vaigat formation is in fact onlapped by a range of units of the maligât formation. this onlap can be observed in long and well-exposed mountainsides, as in the kuugannguaq valley on disko (see fig. 45). in total, the units of the maligât formation resting on the vaigat formation change from the lower rinks dal member (unit 509–511) on disko to the lower to upper rinks dal member (unit 511–514) on nuussuaq (figs 9–11, 15). the older units (505–507) are not known to be in contact with the vaigat formation and it is theoretically possible that they may be contemporaneous with the youngest part of the vaigat formation. atanikerluk formation. within the large sedimentary basin with the assoq lake on eastern disko and southeastern nuussuaq (figs 15, 16), the lower boundary of the maligât formation is marked by basaltic lava flows and hyaloclastites which overlie or are emplaced into mudstones and sandstones of the assoq and umiussat members of the atanikerluk formation (see section on the assoq lake above). the boundary zone is characterised by a range of volcanic morphologies created as the volcanic rocks entered the lake, swamps and fluvial plains 0.5 m mf11_11cmimg_4463akpcut2_c.eps fig. 17. the lower boundary of the maligât formation on south disko between akuarut/ lyngmarksfjeld and navaranaat/apostelfjeld. a basalt lava flow (unit 507) transitional to pillow lava rests directly on an eroded, but not weathered, gneiss surface. 41 mf12_17cmfortune_2_08mod2_c.eps 514 513 507 509 507 506 gneiss gneiss 506 fig. 18. the lower part of the maligât formation on south disko at tuapassuit, resting on gneiss (see fig. 27). the boundary to the gneiss is concealed by talus fans except at the small white circle, which indicates the location of the images in fig. 19. all lava flows are subaerial, and the lowest flow here belongs to unit 506. the plateau top to the left is at 560 m a.s.l. mf13a_8cmakp_1984_10_25udcut2_c.eps 506 a mf13b_8cma1984_10_28cut3_c.eps b 506 fig. 19. deeply weathered surface of the basement gneiss beneath lava flows of the maligât formation. a: saprolite consisting of weathered and oxidised gneiss overlain by a lava flow of unit 506. the sediment is undercut by a small stream. b: kaolinised and disintegrated gneiss (at hammer) overlain by dark grey mudstone. in the centre of the mudstone is a small debris flow with angular, poorly sorted clasts of kaolinised gneiss and mudstone that has a load-deformed base. the mudstone is covered by highly oxidised lateritic soil containing a dark stone (in shadow). the base of the overrunning lava flow of unit 506 is traced. length of hammer 32 cm. tuapassuit, south disko (for location, see fig. 18). 4242 with unconsolidated mud, sand and peat. the units of the maligât formation in contact with the sediments range from the lowest rinks dal member (unit 506) on central disko to the middle rinks dal member (unit 513) on eastern disko and south-eastern nuussuaq. on north-eastern nuussuaq east of the boundary fault, lava flows of the upper niaqussat member (unit 532) are intercalated with sediments assigned to the atanikerluk formation. the detailed contact relations are described in the sections on the individual members and units of the maligât formation. upper boundary the upper boundary of the maligât formation on disko and nuussuaq is erosional. on disko the highest parts of the formation are exposed on western disko where they include three lava flows of the sapernuvik member (unit 540). on north-eastern nuussuaq the uppermost part of the maligât formation comprises basaltic lava flows from the uppermost niaqussat member (unit 532). fig. 20. near-conformable boundary between the vaigat and maligât formations in subaerial facies. thick brown lava flows of maligât formation units 513–514 overlie thin grey picrite flows of the ordlingassoq member of the vaigat formation. the preserved part of the maligât formation is 150 m thick in the central peak. north wall of ataata kuua, nuussuaq. 43 rinks dal member lithostratigraphy of the rinks dal member revised member history. the rinks dal member was informally established by pedersen (1975a) on north-western disko. here the definition is formalised and extended to cover the volcanic areas on both disko and nuussuaq east of the itilli fault. name. after rink dal (formerly spelled rinks dal) on north-western disko. distribution. the rinks dal member extends over the whole of disko and parts of southern and central nuussuaq (fig. 15). the eastern and northern limits are depositional and relatively well defined, whereas the southern and western limits are unknown. type section. the south side of the mountain orpiit qaqqaat at 0–1080 m, central disko. this comprises the thickest and most complete section through the rinks dal member although the uppermost flows are missing (fig. 8, profile 4); fig. 24 (photogrammetrically measured section), fig. 28 (geological map) and figs 25, 31, 63 (photographs). the uppermost flows and the boundary to the nordfjord member are covered by a short profile on the eastern side of the mountain akuliarusersuaq at 990–1045 m, 13 km farther south (fig. 8, profile 3). reference sections. the eastern wall of brededal at 0–920 m, c. 3 km from the south coast of disko (fig. 7, profile 5; south disko section at 95–96 km). the eastern wall of the kuugannguaq valley just south of 70°n, at 760–1600 m, northern disko (fig. 9, profile 5; central disko section at 56–57 km). the southern shoulder of the mountain qinngusaq leading down to the innermost part of kvandalen, at c. 500–1150 m, eastern disko (fig. 10, profile 8; central disko section at 79–83 km). giesecke monument (uppalluk) and the neighbouring mountain (point 1580 m) at c. 850–1550 m, south coast of nuussuaq (fig. 11, profile 7; south nuussuaq section at 57–59 km; figs 74, 75, 94). the mountain point 1888 m west of ataataa kuua, at 1565–1860 m, 7.5 km from the south coast of nuussuaq (fig. 11, profile 2). nunavik, the southern slopes below the mountain point 2000 m, at c. 1300–1650 m, eastern nuussuaq (fig. 12, profile 13; central nuussuaq section at 67–68 km, fig. 95). thickness. the rinks dal member is thickest on western and central disko around the disko gneiss ridge where the lower boundary is usually below exposure level. the exposed thickness here is generally around 1000 m and reaches up to 1460 m. the member is thinner where it engulfs the disko gneiss ridge with elevations up to 750 m a.s.l.. east of the ridge, where both the lower and upper boundaries are exposed, the thickness is 1000 m on southern and central disko, thinning towards the north where the rinks dal member onlaps the vaigat formation, and also thinning towards the east where the member interfingers with contemporaneous fluvial and lacustrine sediments. the member is c. 300 m thick around aqajaruata qaqqaa on easternmost disko, and 800 m thick in the inner part of the kuugannguaq valley on northern disko. on nuussuaq, a maximum thickness of 600 m is attained around giesecke monument, whereas on central and eastern nuussuaq the thickness is reduced to 200–300 m because of the onlap on the vaigat formation and on the eastern gneiss highland. thicknesses on western and northern nuussuaq cannot be estimated because of erosion. lithology. plagioclase-phyric and aphyric tholeiitic basalts of uniform aspect. mainly subaeriel lava flows, but east of the disko gneiss ridge also subaqueous and invasive lava flows and hyaloclastites. the lava flows vary in thickness from less than 5 m to 60 m; most flows are 15–30 m thick. in a few cases, ponded flows are up to 100 m thick. interbasaltic sediments in the subaerial lava succession are mainly thin lateritic top soils on lava flows. subdivisions. the rinks dal member (rdm) is subdivided into 12 chemostratigraphic units which are all informal because they are generally not lithologically distinguishable. however, one large and extensive unit in the middle of the rdm is partly mappable and is shown on the geological maps on a scale of 1:100 000 of disko south of 70°n (unit βi) and south-east nuussuaq (unit mi). this unit is characterised by high contents of iron and titanium and was called the feti unit by larsen & pedersen (1990); it is here named the akuarut unit. a 4444 smaller unit is only mappable east of the disko gneiss ridge where it is in hyaloclastite facies; this unit is here named the skarvefjeld unit (formerly the ‘pahoehoe unit’ of larsen & pedersen 1990) and is shown on the geological maps on a scale of 1:100 000 of eastern and southern disko (unit βfph2). in the following, it has been convenient to assemble the 12 units of the rdm into three parts, the lower, middle and upper rdm, with the akuarut unit constituting the middle rdm. boundaries. the lower boundary is below exposure level on western disko. where exposed, the lower boundary is unconformable. it is also diachronous, younging towards the east and north. the rinks dal member onlaps the disko gneiss ridge on central and south disko, as well as the lavas of the vaigat formation on northern disko and nuussuaq. it invades and interfingers with contemporaneous fluvial and lacustrine sediments on southern and eastern disko and south-eastern nuussuaq. at the upper boundary the rinks dal member is conformably overlain by lava flows of the nordfjord member; there is commonly, but not always, a sediment horizon at the boundary. age. paleocene, 61–60 ma, magnetochron c26r, based on radiometric dating (storey et al. 1998; larsen et al. 2016). correlation. none certain. the relationship to the nûluk member on hareøen and westernmost nuussuaq west of the itilli fault is uncertain. subdivision of the rinks dal member the rinks dal member is the thickest and most extensive member of the maligât formation. it constitutes the major part of the preserved succession and extends over all of disko where it is well exposed along the south coast (south disko section) and along many valley sides, e.g. in the central disko section. it caps the high mountaintops of southern and south-eastern nuussuaq (south nuussuaq section) and is present at high altitudes on northeastern nuussuaq (central nuussuaq section, north nuussuaq section). because of the lithological uniformity of the rinks dal member, its subdivision is based on geochemical variations in many flow-by-flow sampled profiles. in the following, all quoted element concentrations are from analyses recalculated to 100 wt% on a volatile-free basis. the rinks dal member overwhelmingly consists of three-phase-cotectic basalts, i.e. relatively evolved melts in equilibrium with olivine, plagioclase and clinopyroxene. of 1011 analysed samples, 998 have mgo = 4.4–9.2 wt% (average 6.4 wt%); only 13 samples have mgo >10 wt% and most of these contain accumulated olivine. the 998 samples have tio2 = 1.47–4.78 wt% (average 2.73 wt%) and k2o = 0.06–1.05 wt% (average 0.30 wt%). sio2 contents are 46.4–50.7 wt% except for two samples with 51.3–51.4 wt%; crustal contamination is thus very rare. 0 100 200 300 400 0 1 2 3 4 5 6 sy nt he tic su cc es sio n nu m be r tio2 (wt%) 514 513 512 511 507 506506 505505 509 515 518 517 516 akuarut unit skarvefjeld unit fig. 21. the variation in tio2 content with height in the rinks dal member. the vertical axis is a synthetic succession numbered 1–450, comprising 440 analysed samples arranged in stratigraphic order; the number of samples in each unit approximately reflects the correct volume proportions. unit divisions are shown with threedigit unit codes (table 1) on the left. low tio2 contents in some samples in the skarvefjeld unit (511) are due to high mgo contents not seen in this plot. 45 the main subdivision of the rinks dal member is three-fold (fig. 5). the backbone of the subdivision is the akuarut unit, a succession of ti-rich flows in the middle part of the rinks dal member that is easily recognisable in the analysed profiles and is present over all of disko and southern nuussuaq. the akuarut unit conveniently divides the rinks dal member into the three parts shown on the newer geological maps and the geological sections. in detail, the rinks dal member is subdivided into 12 chemostratigraphic units based on oscillations in the tio2 content with height (fig. 21 and table 2). this subdivision forms the basis for lateral correlations over wide areas (figs 7–14). the subdivision by chemical means has some drawbacks. the boundaries between units are ‘artificial’, the more so when the geochemical changes with height are gradual. for instance, the akuarut unit is surrounded by ‘transitional’ flows that have compositions intermediate between the akuarut unit and the older and younger flows (fig. 21). moreover, flows from some units may interdigitate; a good example of this is again the akuarut unit, which on northern disko and southern nuussuaq splits up in two intervals with ‘transitional’ flows between them. lower rinks dal member summary of the main features of the lower rinks dal member • uneven substrate: disko gneiss ridge hills, assoq lake basin east of the ridge and vaigat formation picrite shield rising to the north. substrate to the west unknown. • earliest units (505, 506, 507) are only present on south-central disko. later units (509, 511, 512) spread gradually over all of disko; 511 and 512 also occur on south nuussuaq. • unit 509 is the first unit known to step onto the vaigat formation picrite shield on central disko. • gradual drowning of nearly all the remaining hills of the disko gneiss ridge by lava flows. • gradual eastward filling of the assoq lake basin with hyaloclastites and subaqueous lava flows. several filling episodes can be distinguished. • lava flows at the eastern margins of the lava plateau frequently invasive into fluvial and lake sediments of the assoq member. • individual units are distinguishable chemically but generally not lithologically. parts of units 507 (low-ti unit) and 511 (skarvefjeld unit) are mappable because they form hyaloclastites and thin-bedded pahoehoe flow groups. • unit 511 (skarvefjeld unit) represents an episode with eruption of mg-rich magmas interpreted as due to incomplete mixing with a new picritic magma batch intruded into the deep-seated magma chamber. • unit 512 has increased tio2 and forms a transition to unit 513 (akuarut unit, middle maligât formation). the lower rinks dal member comprises units 505, 506, 507, 509, 511, and 512 (there are no units 508 and 510). the depositional history of this part of the succession is strongly influenced by the uneven substrate and the existence of the large assoq lake basin east of the disko gneiss ridge (fig. 16; dam et al. 2009). the lowermost units (505, 506, 507) are only present on south and central disko (fig. 15). subaerial lava flows spread on the disko gneiss ridge, and subaqueous lava flows, pillow lavas and hyaloclastites filled out the western part of the assoq lake east of the ridge (fig. 16). the two lowest units are below the present sea level or exposure level west of the disko gneiss ridge. they may have been erupted locally, as they most probably could not cross the high-standing ridge, whereas the following unit (507) is also found on the western side of the ridge at killiit/fortunebay (fig. 7, profile 2). the following unit 509 was the first to spread over large areas of disko (fig. 15), onlapping the vaigat formation in the north and extending farther into the subsiding assoq lake basin. units 511 and 512 continued this spreading and infilling pattern until most of disko was a flat subaerial lava plain bordering a fluvial plain with coal swamps on eastern disko and southern nuussuaq. unit code designation tio2 range (wt%) 518 uppermost flows 2.3–4.5 517 1.7–3.4 516 upper low-ti unit 1.5–2.0 515 1.8–3.4 514 upper transition flows 2.4– <3.2 513 akuarut unit ≥ 3.2 512 lower transition flows 2.4– <3.2 511 skarvefjeld unit* 2.06–2.78 † 509 1.8–2.5 507 lower low-ti unit 1.5–1.8 506 1.8–2.6 505 lowest flows 2.9–3.2 *this unit is distinguished by higher tio2 for a given mgo. †except in samples with mgo contents >12 wt%. unit codes 508 and 510 are unused. table 2. subdivision of the rinks dal member by tio2 contents 4646 unit 505 (lowest flows) composition and petrography. the lowest flows have relatively high tio2 contents of 2.9–3.2 wt%. the typical basalt contains scattered microphenocrysts and small glomerocrysts (less than 2 mm in size) of plagioclase and augite. distribution. the lowermost flows have only been sampled in two profiles in a narrow area on the eastern flank of the disko gneiss ridge (fig. 15). at eqalunnguaqqat qaqqaat (fig. 8, profile 2), unit 505 is at least 170 m thick, and c. 10 km to the east at the type locality orpiit qaqqaat (fig. 8, profile 4) it is more than 190 m thick. lithologies. at eqalunnguaqqat qaqqaat the lowermost basalt is a 40 m thick flow overlying conglomerate and sediment developed on the gneiss surface; it is transitional between a subaqueous lava flow and a pillow lava with pillows several metres in size. it is covered by a few small lava flows and by a large, c. 100 m thick ponded lava (f. ulff-møller, unpublished field notes 1978) which is in turn covered by several subaerial lava flows with redoxidised, scoria-rich flow tops. unit 505 can be followed eastwards to orpiit qaqqaat as poorly exposed hyaloclastite (fig. 22) and pillow tongues dipping in an easterly direction. several flow units of hyaloclastite, each thickening eastwards, are pres ent; the lowermost of these is c. 150 m thick and indicates a water depth in the lake basin of this size (fig. 8, profile 4). the base of unit 505 is below sea level near orpiit qaqqaat, but other exposures in the area indicate that the hyaloclastites here overlie lacustrine mudstones of the assoq member of the atanikerluk formation. unit 506 composition and petrography. unit 506 has tio2 contents of 1.8–2.3 wt% with a few values of up to 2.6 wt%. typical basalts are phenocryst-poor with scattered microphenocrysts and phenocrysts of plagioclase and less common augite up to 2 mm in size, and with pseudomorphs after olivine less than 1 mm in size (fig. 23). these basalts are lithologically and chemically similar to those of unit 509 but are separated from them by the distinctive unit 507. distribution. the basalts of unit 506 occupy an elongate, c. 25 km wide, n–s-trending belt confined to the west by the disko gneiss ridge and to the east by lack of exposures (fig. 15). on the disko gneiss ridge the distribution and thickness of the unit is affected by the considerable palaeo-topography (bonow 2005). the unit here comprises just a few flows, but its thickness increases eastwards to up to 225 m on central disko at orpiit qaqqaat (figs 8, 24) where the lavas prograded towards the east and north into the assoq lake basin. lithologies. unit 506 has been sampled and recorded at a few localities on the south coast of disko. a single lava flow overlies saprolite at tuapassuit (fig. 7, profile 3), and about 100 m north of qeqertarsuaq town pillow lavas and hyaloclastites overlie local sands (the main aquifer of the town). at brededal, a c. 40 m thick subaqueous fig. 22. oldest exposed deposits of the maligât formation: hyaloclastites rich in pillow fragments, unit 505, lowest rinks dal member. length of swiss army knife 8 cm. orpiit qaqqaat, south-central disko. 47 lava flow overlies mudstone of the assoq member (fig. 7, profile 5). the main studied occurrence of unit 506 is situated along the inner part of the kuannersuit sulluat fjord on the lower slopes of the mountain orpiit qaqqaat (fig. 8, profile 4; fig. 24) where the unit is moderately well exposed over c. 5 km (fig. 25). it is c. 225 m thick and comprises a number of lava flow fields showing general thickening from west to east as they filled the growing accommodation space in the subsiding lake basin. a wellexposed gully (fig. 25) displays a considerable lithological variation showing that subsidence and basin filling were episodic. here, unit 506 rests on foreset-bedded hyaloclastites of unit 505, and four episodes of deposition of unit 506, as well as two episodes of drowning, can be distinguished (fig. 24). filling episode 1 led to the deposition of several subaqueous lava flows with very well-developed colonnades and hyaloclastite tops (in all c. 60 m thick), followed by a c. 10 m thick flow unit with mega-pillows and pillow tubes. this is capped by a 12 m thick flow unit of thin subaerial pahoehoe lobes with lateritic soil on the top, indicating complete emergence. a first drowning episode is indicated by the renewed occurrence of subaqueous volcanic facies. filling episodes 2 and 3 led to the deposition of two subaqueous lava flows covered by 3–4 m of basalt gravel and conglomerate (episode 2, partial filling), followed by c. 14 m of foreset-bedded hyaloclastite and 10 m of associated thin subaerial pahoehoe flows which constitute one flow field. this is followed by a c. 15 m thick, distinctly subaerial lava flow with an oxidised, scoria-rich flow top, capped by lateritic soil (episode 3, complete filling). a second, lesser drowning episode is indicated by a renewed occurrence of water-influenced lava morphologies. filling episode 4 led to the deposition of a 40–45 m thick lava flow with a regular colonnade and entablature indicating flow into a wet environment. however, the metre-thick, oxidised flow top shows that the basin was filled well above the wet zone. unit 506 is here capped by a flow field of thin subaerial pahoehoe lobes of unit 507. unit 507 (low-ti flows) composition and petrography. unit 507 is a distinctive lithological and chemical marker horizon. it consists of low-titanium basalts with tio2 contents of 1.4–1.8 wt% (fig. 21), rich in phenocrysts of plagioclase and less common augite. the phenocrysts tend to form glomerophyric clusters up to c. 4 mm in size but occasionally up to 10 mm. olivine (pseudomorphs) form scattered microphenocrysts less than 1 mm in size. (fig. 26). in wet facies the basaltic groundmass is very fine-grained, whereas subaerial basalts have well crystallised, almost doleritic groundmass textures. distribution. the basalts of unit 507 occupy a c. 35 km wide, north–south elongate belt confined towards both west and east by being below exposure level (fig. 15). on the disko gneiss ridge, the distribution and thicknesses are still affected by the palaeo-topography. the unit is thickest in its easternmost parts (360 m in the profile north of sorte hak, fig. 9, profile 4) because of infill of the assoq lake basin. the lavas of unit 507 are known from two main areas on south and central disko. they occur in a 17 km long section along the south coast of disko from killiit in the west to blæsedalen in the east; they are shown as units m1 mf17_8cm354727upcut2_c.eps cpx pl ol 5 mm fig. 23. thin section (scanned) of a typical lava from rinks dal member unit 506. basalt with scattered phenocrysts of plagioclase (pl), olivine (ol, altered to dark green clay) and augite (cpx) in a fine-grained matrix. sample ggu 354727, orpiit qaqqaat, southcentral disko. 4848 32 84 90 o rp iit q aq qa at ty pe se ct io n fo r t he r in ks d al m em be r 35 47 04 35 47 33 ? 35 47 19 -2 0 c gl. 32 84 79 20 0 1 2 1. 5 0. 5 2. 5 3 3. 5 ? ? 4 km 40 0 60 0 80 0 10 00 12 00 14 00 m w e fi g. 2 2 u ni t 5 18 u ni t 5 17 u ni t 5 16 u ni t 5 15 u ni t 5 13 ( ak ua ru t u ni t) u ni t 5 12 u ni t 5 11 (s ka rv ef je ld u ni t) u ni t 5 09 –5 11 u nd iff er en tia te d u ni t 5 09 u ni t 5 07 su ba er ial la va fl ow s u ni t 5 07 fo re se t-b ed de d hy alo cla st ite s u ni t 5 06 l ar ge p on de d lav a flo w u ni t 5 06 su ba qu eo us la va fl ow s u ni t 5 06 su ba er ial a nd su ba qu eo us la va to ng ue s la te rit ic so il u ni t 5 05 p illo w la va s a nd h ya lo cla st ite s sa m pl e fa ul t c on glo m er at e c ol on na de fo re se t b ed di ng c gl. 35 47 18 g v0 3_ 03 _0 47 _l m l. ep s fi g. 2 4. t he ty pe se ct io n fo r t he r in ks d al m em be r w ith th e t hi ck es t a nd m os t c om pl et e s ec tio n th ro ug h th e m em be r. se e m ap in f ig . 2 8 an d ph ot og ra ph s i n fi gs 2 5, 3 1, 6 3. t he in di ca te d sa m pl es a re us ed to co ns tr uc t t he p ro fil e s ho w n in f ig s 8 , 9 . p ho to gr am m et ric al ly m ea su re d se ct io n, so ut h sid e o f t he m ou nt ai n o rp iit q aq qa at , s ou th -c en tr al d isk o. 49 and m2 in the south disko section at 65–82 km (see fig. 36) and as parts of the units βf1 and βf1u on the geological map sheet uiffaq (fig. 27). on central disko, unit 507 occurs on both sides of the inner part of the fjord kuannersuit sulluat and on both sides of the valley kuannersuit kuussuat and its side valley sorte hak. it is mapped as unit βfph1 on the geological map sheet pingu (pedersen et al. 2001; fig. 28). unit 507 has a total thickness of 360 m north of sorte hak (fig. 9, profile 3) of which the lower 295 m is in subaqueous facies which makes it one of the thickest subaqueous successions in the maligât formation. south coast of disko. the disko gneiss ridge is exposed and reaches heights up to c. 350 m at the south coast of disko between killiit in the west and rødeelv in the east (figs 6, 27). the relationship between the gneiss topography, subaerial lava flows and subaqueous hyaloclastite breccias is discussed below at four localities (south disko section at 67–81 km, see fig. 36). at killiit at least four subaerial lava flows of unit 507 onlap the western flank of the disko gneiss ridge, but the contact is not exposed. the flow packet is 115 m thick and the flows vary in thickness between 15 and 50 m (fig. 7, profile 2). the uppermost flow is overlain by c. 30 cm of reddish lateritic soil. about 4–5 km farther to the east, north-east of tini, 30–35 m of unit 507 are exposed in a gully. the two localities are separated by a gneiss hill that rises up to 200 m above the level of the lavas of unit 507 (see also fig. 36 at 69–72 km), indicating that the flows near tini must have followed an unexposed palaeovalley. the succession near tini starts with a strongly columnar-jointed subaqueous lava flow covered by a hyaloclastite composed of decimetre-sized pillows and matrix glass shards. the hyaloclastite contains decimetre-sized clasts of claystone and lateritic soil and mf19_17cmlml_1991_14_16mod_c.eps cgl. 513 507 506 505 fig. 25. the western part of the type section for the rinks dal member; unit codes are indicated. cgl: conglomerate. the photo corresponds to fig. 24 at 0.9–2 km; unit 506 (subaqueous and ponded flows) is 230 m thick. south side of the orpiit qaqqaat mountain, south-central disko. 5050 centimetre-sized clasts of gneiss, all of which have been stripped off a weathered gneiss surface. at the contact between the subaqueous flow and the hyaloclastite there are patches of red-oxidised scoria and up to 15 cm of lateritic soil. the top of the hyaloclastite is covered by several decimetres of reddish-brown soil of subaerial origin. this is in turn covered by a compound pahoehoe flow indicating subaerial conditions. east of tini, at tuapassuit, unit 507 forms a 115 m thick succession of six subaerial lava flows above a 55 cm thick horizon of sandstone and siltstone with plant remains overlying unit 506 (fig. 7, profile 3). most of unit 507 must bank up against the disko gneiss ridge northwest of qeqertarsuaq town because it is not present on the top of the ridge (see fig. 36 at 76–77.5 km). a single flow of unit 507 overlies the gneiss ridge in the canyon between apostelfjeld and lyngmarksfjeld (fig. 29). mf20_8cm332960upcut_c.eps cpx pl cpx ol 5 mm fig. 26. thin section (scanned) of a typical lava from rinks dal member unit 507. highly porphyritic basalt with many phenocrysts of plagioclase (pl), olivine (ol, altered to brown clay and white zeolite) and augite (cpx) in a medium-grained matrix. sample ggu 332960, sorte hak, central disko. ng 1f 2f u u u u ef u1f 1f ng 5.1 5.1 5.0 1 1 1 5.1 1 2.5 km 54°45’ 54°30’ gv01_02_155_lml 60°15’ βf2 βf1 βf1u oqaatsunnguit qaqqaat qausrotrolroq eq g qatreqeq nełdłk tiillik yabenutrof turauka dlejfskramgnyl taanaravan dlejfletsopa ku us su aq r ød ee lv qailaqqaq neggikdu tiusuraaq sneksrot init 6 qaarusuit torskenæs qeqertaq kødøen killiit fortunebay qaqqaliaq udkiggen qeqertarsuaq godhavn fe gn gn tini tuapass uit βf1 609 fig. 69 figs 18,19 figs 17, 29 fig. 30 fig. 35 fig. 27. excerpt from the 1:100 000 scale geological map sheet 69 v.1 syd uiffaq, showing the lava succession of the rinks dal member covering the hilly relief on the disko gneiss ridge. the brown unit comprises the fe-ti-rich lava flows of the akuarut unit (513). annotation as on the original map, see fig. 28. the native-iron-bearing killiit dyke is shown as red lines marked fe. the locations of some figures are indicated. 51 this flow has a brecciated lower part and a thin top zone with oxidised scoria and traces of sediment. between tuapassuit and qeqertarsuaq the disko gneiss ridge is c. 200 m high. on its eastern flank, towards the outflow of rødeelv, unit 507 reappears. subaqueous lava flows of this unit are exposed for about 2 km along a rocky beach and a coastal cliff. here a varied succession of large pillow lobes, subaqueous basalt breccias and fine-clastic hyaloclastites formed when a number of subaerial pahoehoe lava lobes entered the assoq lake basin. the elongate pillow lobes (fig. 30) show a general flow direction from the north-west. parts of the lobes have cavities now filled by zeolites. the breccias are composed of pillows with thin glassy rims in a coarse matrix of brownish-weathering palagonitised glass. 2.5 km 53°15’ 69°45’ 53°30’ 1.5 1 0.5 0.5 1.5 1.5 1.5 1 0.5 0.50.5 0.5 1 1.51 1 1 1 2 1075 0.5 0.5 1.5 1.5 2 1.51 1149 1 948 1210 βf3 βi βi βi βf1 βf1 βf1 βf1u βf1u βf2 βf1u βf2 βf2 pmu pmu t is is βf1 βf1 βf1 βfph1 βfph1 βfph1 βfph1 βfph1 1282 1 1450 blomsterdalen orpiit qaqqaat ku an ne rs ui t k uu ss ua t kuannersuit is is orpiit ang uuja artuutit pullanni sorte hak gv01_02_156_lml fig. 63 fig. 24 fig. 31 fig. 25 fig. 33 fig. 32 fig. 28. excerpt from the 1:100 000 scale geological map sheet 69 v.2 nord pingu, showing almost the full stratigraphy of the rinks dal member. annotation as on the original map. this corresponds to the more detailed lithological units in this work as follows: βf1u is units 505–506, βfph1is unit 507, βf1 is units 509–512, βi is unit 513 and βf2 is units 514–518. βf3 is the overlying lava flows of the nordfjord member, and pmu is small exposures of the underlying mudstones of the assoq member. the line of view of the type section (fig. 24) is shown, as well as the locations of some figures. 5252 central disko. the basalts of unit 507 extend from the south coast of disko northwards to central disko, but they have not been studied or sampled between the south coast and the inner part of kuannersuit sulluat (easternmost kangerluk). unit 507 is present in the profile at eqalunnguaqqat qaqqaat (fig. 8, profile 2) as two more than 30 m thick subaerial lava flows overlying unit 506. in the gully on the south slope of orpiit qaqqat, unit 507 consists of a small flow field of pahoehoe lobes about 12 m thick covered by a c. 12 m thick lava flow (fig. 8, profile 4). eastwards from there unit 507 rapidly increases in thickness to more than 100 m just 2 km farther to the east, where it changes facies to form foreset-bedded hyaloclastite infill into the assoq lake. the foresets show infilling from the west and north-west, but a very dense joint system striking nw to nnw locally obscures the breccia structure when seen at a distance (fig. 31). the infill into the lake occurred over a more than 20 km long, north–south-extending front that prograded more than 5 km eastwards, reducing the western lake area by more than 100 km2 (figs 15, 16). mf23_8cmimg_4467akpcut_c.eps cpx pl top zone breccia gneiss fig. 29. lava flow of unit 507 overlying the disko gneiss ridge in the canyon between apostelfjeld and lyngmarksfjeld (south disko section at 77.6 km, fig. 36). the flow is c. 30 m thick and in transitional facies between subaqueous and subaerial: it has a brecciated lower part, its interior is a coarse entablature, and there is a thin rubbly top zone with oxidised scoria. mf24_11cmimg_0618akp_c.eps 1 m fig. 30. metre-long pillow lobe from a pahoehoe flow of unit 507 that flowed into the assoq lake. the lobe has a near-tubular central cavity filled by zeolites. near the outflow of rødeelv east of qeqertarsuaq town, south disko. 53 up to 10 m of black mudstone of the assoq member from the lake floor is exposed at the base of the volcanic rocks at a few localities (unit pmu on the map sheet pingu, fig. 28). in the profile 5 km north of sorte hak, the contact between mudstone and infilling hyaloclastites has produced a chaotic mixture like very coarse peperite where pillow tongues invaded unconsolidated mud (fig. 32). on both walls of the kuannersuit kuussuat valley and its side valley sorte hak (fig. 28), unit 507 consists of foreset-bedded hyaloclastites and overlying subaerial pahoehoe lava flows. excellent exposures in the eastern wall of kuannersuit kuussuat north of sorte hak formerly showed the presence of several distinct hyaloclastite beds which cannot have been formed during a single eruptive event. these exposures were covered in 1995 by a surging glacier lobe. the highest foresets produced in a single filling event, which record the water depth in the assoq lake at that time, are around 100–150 m high and occur in the northern wall of sorte hak. the subaerial pahoehoe lava flows comprise 10 to 15 individual lava lobes in vertical section and represent several flow fields. the uppermost lava flow of unit 507 north of sorte hak is thick (35–40 m) and massive and has a regular colonnade indicating that it solidified in a moist environment (fig. 33). mf25_17cmlml_1991_14_10mod_c.eps 507 hy 506 sla 509 507 la 511 fig. 31. middle part of the type section of the rinks dal member. the photo corresponds to fig. 24 at 2.8–3.5 km. unit 506 is in subaqueous lava facies (506 sla). a subaerial succession of thin pahoehoe flows of unit 507 (507 la) entered the assoq lake from the west and formed a thick horizon of massive foreset-bedded hyaloclastites (507 hy) that prograded more than 5 km into the lake. the foresets are obscured by a very pronounced vertical joint system. the overlying flows are all subaerial. the combined unit 507 (flows and hyaloclastites) is 150 m thick. south side of the orpiit qaqqaat mountain, south-central disko. 5454 fig. 32. lava lobes and pillows of a subaqueous lava flow invading black mudstones of the assoq member. the hammer is 32 cm long. valley bottom 5 km north of sorte hak (fig. 28; fig. 9, base of profile 4); the locality was covered by ice after a major glacier surge in 1995–1999. mf27_17cmlml_1991_11_36cut_c.eps 507 la 507 la 507 hy sediment fig. 33. lithological variations in unit 507. the unit comprises hyaloclastites (507 hy, the foresets dip away from the viewer), a thin sediment horizon, subaerial pahoehoe flows (507 la) and a thick, massive flow with a well-developed colonnade (also 507 la). the thick flow is 35–40 m thick. kuannersuit kuussuat north of sorte hak, central disko. 55 unit 509 composition and petrography. unit 509 is composed of basalts with tio2 contents of 1.8–2.5 wt% (fig. 21). they contain scattered plagioclase and augite glomerocrysts between 1 and 3 mm in size and pseudomorphs up to 1 mm large after olivine crystals (fig. 34a) in a fine-grained to glassy groundmass. some regional variations are present. the subaqueous lava flows along the south coast of disko from ippik (east of skarvefjeld) to tuapaat qaqqaat have very uniform compositions with tio2 contents of 2.0–2.2 wt% and mg-numbers of 48– 55; these flows were called the ippik unit by larsen et al. (2006). on central and northern disko, several lava flows of unit 509 have higher mg-numbers (57–62), corresponding to mgo contents of 7.7–8.8 wt%; these flows may be aphyric to slightly olivine-microphyric (fig. 34b). other northern flows are very similar to those on southern disko. distribution. unit 509 is present over large parts of central and southern disko (fig. 15); to the north, it reaches qullissat and asuutaa but peters out before orlingasoq (fig. 10). it also peters out on eastern disko and is not present east of tuapaat qaqqaat (fig. 7), laksedal (fig. 8) and inussuk (fig. 10). it is present on western disko around inner nordfjord (fig. 13) and is assumed to be present beneath sea or exposure level in large areas west of the disko gneiss ridge. excluding onlap settings, thicknesses are between 50 m and 240 m. unit 509 almost, but not completely, covered the disko gneiss ridge. the unit is in subaerial facies on the gneiss ridge and west of it. east of the gneiss ridge the lavas plunged into the assoq lake and continued as coherent subaqueous lava flows. the unit prograded far into the assoq lake where the flows burrowed into lacustrine and fluvial sediments and continued as invasive lava flows that resemble sills (south disko section). lithologies. in subaerial facies, unit 509 comprises 2–10 lava flows per profile, commonly around five flows. the combined thickness of the flows may exceed 150 m. individual flows vary in thickness from less than 5 m to more than 50 m, typical flows being 10–30 m thick. the greyish weathering flows are of blocky pahoehoe type (fig. 5 mm mf28a_8cm332986upcut_c.eps a 5 mm mf28b_8cm332881upcut_c.eps b fig. 34. thin sections (scanned) of typical lavas from rinks dal member unit 509. a: basalt with numerous glomerocrysts of plagioclase (white), augite (grey) and olivine (dark brown, altered to clay) and smaller single plagioclase and olivine phenocrysts in a fine-grained matrix. sample ggu 332986, daugaard-jensen dal, south disko. b: nearly aphyric basalt with a few microphenocrysts of olivine (dark brown, altered to clay) and plagioclase. sample ggu 332881, kuugannguaq, northern disko. 5656 35; kilburn 2000) and are visually indistinguishable from most other subaerial flows of the rinks dal member. the flows onlap the flank of the disko gneiss ridge at killiit (fig. 7) and west of tini but overflowed the ridge farther to the east and north (fig. 36). the lava flows entered the assoq lake in an easterly direction along a palaeoshore that extended for more than 50 km from the south coast of disko at blæsedalen and northwards, where much is presently below exposure level. the infill of unit 509 into the lake must have reduced the lake area by more than 1750 km2. a northern lake or wetland, which may have been continuous or contemporaneous with the assoq lake, existed on north-eastern disko between kuugannguaq and qullissat (fig. 15, see also fig. 48). the large and relatively thick lava flows of unit 509 that entered the assoq lake maintained their coherence, also at subaqueous conditions (e.g. fig. 56). this is in contrast to the small pahoehoe lobes in the flow fields of unit 507 (as well as the overlying unit 511) which were transformed into pillow lavas and foreset-bedded hyaloclastites. a factor contributing to this difference is probably higher magma flow rates in the large flows of unit 509. south coast of disko: infilling of the southern part of the assoq lake the lake floor. mudstones of the assoq member underlie the volcanic units in the assoq lake over large areas on south and south-east disko (dam et al. 2009, fig. 124), but there are very few exposures due to extensive landslides and solifluction. the contact relations between the subaqueous lavas of unit 509 and the mudstones are therefore only seen at few localities. the westernmost mudstone crops out beneath skarvefjeld in a small landslide, but the main exposures are farther to the east at assoq (larsen et al. 2006, fig. 5; dam et al. 2009, figs 138, 139; fig. 7, profile 6), niuluut (larsen et al. 2006, fig. 13), and tuapaat qaqqaat (dam et al. 2009, fig. 140; fig. 7, profile 8). the depth of the assoq lake when unit 509 was emplaced was estimated to 200 ± 100 m by larsen et al. (2006). subaqueous lava flows. the south coast of disko contains remarkable exposures of strongly columnar-jointed basalts with brecciated tops interpreted to be subaqueous lava flows (heinesen 1987; larsen et al. 2006). the subaqueous lava flows extend for more than 25 km from blæsedalen in the west to east of siniffik (south disko section at 82–107 km, unit m3) and their total thickness is at least 200 m, but over large distances east of assoq the exposures are poor and fragmented by landslides. the best and most coherent exposures are found between blæsedalen and brededal. the individual subaqueous flows may well exceed 10 km in length; however, the chemical and petrographical homogeneity of unit 509 make correlation across valleys, landslides and talus virtually impossible. several flows have been mapped continuously for 2–3 km along cliff exposures. the thickness of individual flows varies from 5 m to more than 100 m due to local ponding. in typical cliff sections unit 509 consists of 4–5 subaqueous flows. mf29_8cmimg_4470akpmod_c.eps base top fig. 35. typical, c. 30 m thick subaerial lava flow of blocky pahoehoe type with a massive, crudely jointed central part and an oxidised rubbly top. traces of red lateritic soil occur at the top and base. a patch of brick-red laterite is worked into the flow near its base. rinks dal member unit 509, lyngmarksfjeld, south disko. 57 despite their compositional homogeneity, the flows seem to be well separated in time because there are commonly small patches of sediment between them. just east of blæsedalen hyaloclastite and breccia of unit 507 form an e-dipping palaeoslope which is overlain by a thick flow of unit 509 with a regular colonnade at the base and a thick entablature with several tiers; the top is not exposed. there are traces of sediment in small basins at the base of the flow. at sea level around kuannit the spectacular columnar-jointed masses of basalt form a local tourist attraction. in the cliff beneath skarvefjeld the lowermost flow extends from sea level to c. 60 m and has a wavy top surface on which there are accumulations of a few centimetres to decimetres of mudstone in the local depressions (see also fig. 56). a typical subaqueous lava flow such as seen at ippik (fig. 37) has at its base a few metres thick regular colongv03_03_051_lml 7069 75 km7473 76 77 78 79 km 200 400 600 800 m 200 600 ? 800 m 200 400 600 800 m 71 72 km m4 m3t2b m2 gneiss gn fe fe fe gn m7.3 m7.3 m8 m7 m7 m4 m4 m8.29 m8.29 m8 m8 δc m7 m7 m4 m2 cr m7 m7 m2 m8.29 m7.3 ? tuapassuit killiit dyke tini gneiss oqaatsunnguit qaqqaat qaarusuit torskenæs qeqertarsuaq godhavn arktisk station akuarut lyngmarksfjeld navaranaat apostelfjeld 515 516 515 514 513 514 513 509 509 507 507 507 w e w e w e figs 18,19 fig. 30 figs 17, 29 fig. 36. photogrammetrically measured section along the coastal cliffs on south disko between killiit and arktisk station near qeqertarsuaq. blue three-digit numbers are lithological codes; other annotations as in the original. the brown lava succession is the akuarut unit (513). some marker flows are coloured and numbered individually. the locations of some figures are indicated. excerpt from the south disko section (pedersen et al. 2003). 5858 nade which changes abruptly into a 10 m to more than 20 m thick entablature from which lobes may extend upwards into a 5–10 m thick, thoroughly brecciated top zone. within the top zone itself, and in small depressions on the surface, there may be pockets of volcaniclastic sand and gravel (fig. 38) that may contain layers of mudstone with plant remains. there is no red-oxidation at the flow top or between flows. the top zone itself is disturbed by the overlying subaqueous flow. disturbances of the substrate by bulldozing and loading are especially pronounced where a subaqueous lava flow has run onto unconsolidated mudstone (fig. 39). subaqueous rootless cones. at niuluut about 20 km from the lake shore, a rootless cone field developed when the first subaqueous lava that flowed onto the lake floor was locally brecciated throughout, providing outlet points for hydroclastic explosions of the pressurised and heated water in the underlying mud (larsen et al. 2006). the cliff section cuts through the centres of two cones (fig. 40; also shown on a scale of 1:2500 in the south disko secmf31_8cmlml_1981_05_28modcut_c.eps mf32_8cmlml_1981_05_26cut_c.eps fig. 37. typical subaqueous lava flow of unit 509. the basal zone is a regular colonnade a few metres thick, whereas the major part of the flow is in entablature facies with smaller, curved columns that commonly form swirls. the top zone (not seen) is thoroughly brecciated. the part of the flow visible in the photo is c. 35 m thick. ippik, south coast of disko. fig. 38. fine-grained volcaniclastic sediment between two subaqueous lava flows (the lower flow is just beneath the outcrop). the lower part of the sediment is normally graded with a weak lamination; the upper part is homogeneous. note the absence of red oxidation colours. the basal zone of the overlying flow is thoroughly brecciated. ippik, south coast of disko. 59 tion), and others are probably present both seawards and landwards. the cones are small mounds about 25 m high and 100–200 m wide, composed of normally graded and well-bedded volcaniclastic material intimately mingled with mudstone in up to metre-sized clasts, frequently with plant fragments. in both cones, the bedding dips away from a central ‘chimney’ that is considered to be the outlet point for the explosions and later filled with shaly sediment. large pieces of apparent driftwood indicate the proximity of emerged and forested land. detailed descriptions and interpretations are found in larsen et al. (2006). mf33_17cma871231bennymod_c.eps mudstone 509 fig. 39. a large subaqueous lava flow of unit 509 that has filled a depression within unconsolidated mudstones of the assoq member and deformed and probably partly replaced the mud. the colonnade columns are of different heights but always perpendicular to the lower boundary of the flow. person for scale. the locality is also seen in fig. 40. niuluut, south coast of disko. m m m m lava tongues chimney mf34_fig7_lml-rgb._c_vers 2.eps fig. 40. section through a subaqueous rootless cone. the flank of an adjacent cone is seen to the right. see text and larsen et al. (2006) for descriptions. the cone was covered with mud (m) that also filled the central explosion site (chimney). a thick lava flow (white line traces its base) subsequently overran the cones; thin lava tongues invaded the mud overlying the cone flank to the left, and the massive lava subsided into the mud between the cones (fig. 39). height of cliff section c. 25 m. niuluut, south coast of disko. 6060 115 116 117 km m3 m3 k1b m7 m8 200 400 600 800 m m7 t2b t2a t2a tuapaat qaqqaat nuuk ls ls ls sm ls ls ls e e cgl in in in in in ls gv03_03_052_lml 513 509 not exposed wsw ene fig. 73 fig. 72 fig. 41. lava flows of rinks dal member units 509 (blue) and 513 (brown) invasive (in) into paleocene mudstones (t2a) and sandstones (t2b) of the assoq member of the atanikerluk formation. cgl: conglomerate. sm: sill. pale yellow k1b: cretaceous sandstone. e: entablature. ls: landslipped exposures. the locations of some figures are indicated. tuapaat qaqqaat, south coast of disko, excerpt from the south disko section (pedersen et al. 2003). mf36_17cmimg_8486evsmod_c.eps subaq 1 inv 3 m m m m inv 2 inv 1 fig. 42. lava flows of rinks dal member unit 509 invasive into dark brown mudstones (m) of the assoq member at its type locality. there are three successive, invasive, sill-like flows (inv 1 to inv 3) and a thick subaqueous flow (subaq 1). note the colonnades at both base and top of inv 2. the exposure is about 100 m high and is slightly slipped and rotated. assoq just east of brededal, south coast of disko. photo: erik vest sørensen. 61 invasive lava flows. a number of basaltic sill-like bodies crop out over a distance of more than 23 km between assoq and tuapaat qaqqaat (south disko section at 95– 118 km (fig. 41); in small gullies along the coast there are sometimes fragmentary exposures of mudstone beneath, between and above them. the sills are chemically and petrographically identical to the subaerial and subaqueous lava flows of unit 509. the complete absence of any feeder bodies led larsen & pedersen (1990) to conclude that the sills are invasive lava flows formed when subaqueous flows invaded unconsolidated assoq member sediments and continued their eastward flow as sills. the sill-like invasive lava flows may show planar, concordant lower and upper contacts to the mudstone, or they may form more irregular bodies. at the upper contacts, there may occasionally be pillow bodies, and small apophyses may extend from the lavas to form auto-brecciated hyaloclastite. the ‘sills’ may show well-developed upper and lower colonnades and, in the central part, an entablature zone of small, irregular columns. the best exposures are in the type section of the assoq member in a landslide at assoq just east of brededal (fig. 42; dam et al. 2009, fig. 138 and 139; larsen et al 2006, fig. 6). here three to four successive invasive lava flows are embedded in mudstone and underlying sandstone (fig. 7, profile 6). they are covered by 3–4 subaqueous lava flows, which are broken up into several smaller landslides and are not seen in continuity (fig. 42). some samples of the mudstones between the flows contain dinoflagellate cysts (piasecki et al. 1992; h. nøhr-hansen, personal communication 2017). on the south coast of disko invasive lava flows are relatively well preserved along a w–e oriented ridge extending from tuapaat qaqqaat to tuapaat at the coast (south disko section at 116 to 118 km (fig. 41; fig. 7, profile 8). here both the lower and upper contacts between mudstone of the assoq member and four invasive lava flows of unit 509 are preserved. they are overlain by sand with at least two invasive flows of unit 513. the upper boundary of unit 509 is poorly exposed between assoq and tuapaat qaqqaat but seems to have been entirely beneath the surface of the assoq lake (fig. 7). to the west, beneath skarvefjeld, subaqueous lava flows of unit 509 with traces of mudstone on top are covered by up to 160 m of regularly foreset-bedded hyaloclastites of unit 511 (see also fig. 56). heinesen (1987) provided evidence of semi-continuous subsidence of the assoq basin in the western part of skarvefjeld in the time interval of the eruption of the overlying lava plateau; if subsidence also took place during the emplacement of the subaqueous parts the 160 m height of the foresets is a maximum depth for the lake overlying unit 509. daugaard-jensen dal. complete filling of the central assoq lake took place about 30 km farther to the north in daugaard-jensen dal (fig. 8, profile 5) where 210 m of subaqueous lavas and a single subaerial lava flow of unit 509 can be observed in the northern wall of the valley (fig. 43). there are several typical subaqueous lava flows with well-preserved, thick top zones of pillow breccia with metre-sized pillows and pillow tubes. one of the top zones encloses up to metre-sized boulders of basaltic lava, which are talus fragments from the palaeoshore carried into the lake by the later lava flow (fig. 44). the uppermost subaqueous flow is covered by a 35–40 m thick lava flow that has a well-developed colonnade and entablature zone and a thick scoriaceous top zone which is slightly high-temperature oxidised and obviously subaerial. the northern basin in the kuugannguaq–qullissat area lava flows and breccias of unit 509 form the basal part of the maligât formation over large areas on north-eastern disko between the inner kuugannguaq valley and the vaigat coast around qullissat (see also fig. 48). here unit 509 onlaps the vaigat formation, and the volcanic rocks display a considerable range of morphologies depending on the nature of the environment in which they were emplaced. in large areas unit 509 is concealed below younger basalts or glaciers, scree or landslides. however, semi-continuous exposures of the succession occur in three, more than 10 km long, subparallel sections oriented approximately nw–se. two sections form the walls of the inner part of the kuugannguaq valley and are only 2–3 km apart, whereas the third section comprises the coastal slopes along the vaigat strait 15 km north-east of kuugannguaq. kuugannguaq valley. on both walls of the inner part of the kuugannguaq valley (map sheets qutdligssat and pingu), lava flows of unit 509 are seen onlapping the gently se-sloping top surface of the vaigat formation. the geology of both walls has been compiled from long photogrammetric panels, and the succession has been sampled in a profile on the north-eastern wall where unit 509 is 240 m thick (fig. 9, profile 5); the north-eastern wall is shown in the central disko section at 51.5–60.5 km where the sample profile is situated at 55–56 km. the 6262 distances to the nearest investigated exposures of unit 509 to the south and south-east are around 25 km. in kuugannguaq the lavas of unit 509 encountered waterlogged or water-covered areas which were either an extension of the assoq lake or a separate wetland. the volcanic morphologies record two phases of interaction between the inflowing magmas and water in the subsiding basin. the first phase is seen in both valley walls but is best exposed in the south-western wall. here the basal part of unit 509 comprises a group of several strongly brecciated lava flows up to 40 m thick, which can be followed for about 5 km towards north-west, petering out to 0 m by onlap onto the vaigat formation. these flows are overlain by a few entirely subaerial lava flows, which step directly onto the vaigat formation for 5–6 km farther towards north-west (fig. 45). there are three such flows in the south-west wall and two thinner ones in the north-east wall (fig. 46), indicating that the flows arrived from a south-westerly direction. the flows were probably deflected by the se-sloping surface of the vaigat formation and channelled along this. continued basin subsidence led to renewed accumulation of water in the area. a second water-influenced phase is indicated by a renewed occurrence of breccimf37ab_17cmdau1_2_06mod3_c.eps 509 sla 509 sla 509 511 513 514 512 509 sla fig. 43. section illustrating the filling of the western assoq lake by unit 509. numbers are unit codes. unit 509 comprises 210 m of subaqueous lava flows (509 sla) and a single subaerial flow. the overlying units are all in subaerial lava facies. north side of daugaard-jensen dal, central disko; sampled profile in fig. 8, profile 5. 63 ated flows and flows with distinct development of colonnades and entablature zones, followed by a huge ponded lava flow with a large, regular colonnade (figs 45, 46). this lava packet is well exposed on both valley walls and is thicker in the north-east wall than in the south-west wall, suggesting a somewhat deeper basin to the northeast, filled during continued volcanic inflow from the south-west. these lavas extended 5–10 km farther to the north-west than those of the first phase. the large ponded lava flow can be followed for at least 9–10 km on both valley sides and locally attains a thickness of 70–80 m. it covers more than 50 km2 and has a volume well above 2–3 km3. towards the south-east it seems to consist of several separate subaqueous lava tongues. towards the north-west it forms a single distinct flow with load structures towards the underlying volcaniclastic sediment (fig. 46); its colonnade becomes thicker, continuous, and extremely regular, whereas its top gradually develops scoria and becomes high-temperature oxidised. farther towards the north and north-west the flow thins against the sloping surface of the vaigat formation and the colonnade becomes coarse and less regular until the flow has the appearance of an entirely subaerial lava flow (fig. 47). this flow provides the best example from the maligât formation of the morphological variations of a basaltic lava flow emplaced under changing wet and dry conditions. qullissat coastal section. the exposures along the vaigat coast comprise a c. 11 km long, se–nw-oriented section extending from just north of qullissaaqqat (gamle qullissat) in the south-east to killerpaat qaqqarsuat northwest of qullissat, shown in fig. 48. in the north (fig. 48 at 7.3– >11 km), subaerial picritic lava flows of the vaigat formation form a gently se-sloping surface, similarly to the setting in the inner part of kuugannguaq. at 6.2 to 7.3 km, the vaigat formation terminates in c. 200 m high fans of bedded hyaloclastite deposits. these are covered by lacustrine mudstone except where the sediments have been removed by a glacier. the mudstone is tentatively referred to the assoq member on the basis of its mineralogy, indicating that the assoq lake, or a contemporaneous lake, extended into this area. maligât formation lavas of unit 509 overlie sediments and the vaigat formation throughout the section, but to the south, the boundary and the lower part are poorly exposed because of talus and landslides (fig. 48 at 2–6.5 km). unit 509 is up to 150 m thick in the qullissat section and composed of one to four lava flows, most of which are strongly affected by emplacement over a wet or fully water-covered surface (fig. 10, profiles 4, 5). in the mountain wall beneath killarpaat qaqqarsuat three voluminous lava flows are exposed which have very prominent dark grey entablature zones and olive grey brecciated tops (fig. 49). the top zone of the lowermost flow is entirely a subaqueous pillow breccia, whereas the second flow top is dominated by pillow breccia with metre-sized pillows and also contains patches of high-temperature air-oxidised scoria, showing that this flow for a time completely displaced the water. the third flow also has a top of partly glassy breccia and partly oxidised scoria and was emplaced after return of the water to this part of the mf38_8cmlml_1991_12_02modcut2_c.eps 509 sla 1 m fig. 44. top zone of a subaqueous lava flow of unit 509 enclosing a large boulder of subaerial basalt lava as well as the usual pillows and pillow fragments. the lava boulders are older talus fragments from the palaeoshore carried into the lake by the lava flow. north side of daugaard-jensen dal, central disko. 6464 mf39_17cmakp_1991_12_05modcut_c.eps 509 la 450 la fig. 45. lava flows of maligât formation unit 509 onlapping the gently sloping surface of the vaigat formation (450 la). the three lowest flows of unit 509 are subaerial, the following dark flow is brecciated, and the overlying thick flow is ponded. height of exposure c. 400 m. south-western wall of the inner kuugannguaq valley c. 45 km from the mouth of the valley. mf40_11cmlml_1988_08_20udcut_c.eps 509 sla 509 la soil 509 la picrite 450 la volcaniclastic sediment fig. 46. subaerial picrite flows of the vaigat formation unit 450 overlain by basalt flows of the maligât formation unit 509. there are two subaerial flows with a reddish soil horizon between them, a black volcaniclastic sediment and a thick subaqueous flow (509 sla) with irregular base because of loading into the soft sediment. height of field of view c. 100 m. north-eastern wall of the inner kuugannguaq valley c. 45 km from the mouth of the valley. 65 subsiding basin. farther to the north-west, just north of fig. 48, only one of these flows has well developed colonnade and entablature; it has an entirely oxidised and scoriaceous top and marks the final infilling of the wetland in this part of disko. five kilometres farther to the north-west, at asuutaa, unit 509 consists of one or two entirely subaerial lava flows with a total thickness of less than 40 m (fig. 10, profile 2). in the central part of the qullissat section, where vaigat formation hyaloclastites are covered by mudstones and sandstones of the atanikerluk formation, a columnar-jointed sill is intruded into the sediments (fig. 48); see also pedersen et al. (2017, fig. 162). the sill has a composition typical for unit 509 and is interpreted as an invasive lava flow. its continuation into the normal lava package is obscured by landslides. in the south-eastern part of the qullissat section, at inussuk, unit 509 consists of three or four entirely subaqueous lava flows with glassy brecciated top zones. in one of the flows, the top zone is a 10 m thick stratified hyaloclastite. the lower part of unit 509 is here covered by landslides. as in kuugannguaq, the basalts of unit 509 reached the qullissat area from the south-west and onlapped the gently se-sloping top surface of the vaigat formation. unit 511 (skarvefjeld unit) composition. the skarvefjeld unit represents a shift in chemical composition. the defining feature is that, for a given mg-number (atomic 100mg/(mg + fe2+)), the lavas of the skarvefjeld unit have significantly higher tio2 contents than the lavas of the earlier units, so much that in a plot of tio2 versus mg-number there is little or no overlap between the skarvefjeld unit and the earlier units (see also fig. 97). the lavas of the skarvefjeld unit have tio2 contents of 1.9–2.8 wt%; most of them have mgo contents of 6–9 wt% but the unit also includes magnesian rocks with mgo contents of 10–25 wt%; of the 13 samples from the rinks dal member with mgo >10 wt%, 11 are from the skarvefjeld unit. it is sometimes difficult to chemically distinguish the most evolved lavas of unit 511 from unit 512. this is understandable in view of the interpretation by larsen & mf41_17cmakp_1988_10_23mod_c.eps 450 la 450 hy 509 fig. 47. picrite hyaloclastites (450 hy) and subaerial lava flows (450 la) of the vaigat formation overlain by entirely subaerial flows of the maligât formation unit 509. the lowest flow, thinning towards the north (left), is the same as the thick ponded flow shown in figs 45, 46. height of field of view c. 500 m. north-eastern wall of the inner kuugannguaq valley c. 40 km from the mouth of the valley. 6666 pedersen (2009) that the skarvefjeld unit is the result of incomplete mixing of the evolved magma residing in the deep-seated magma chamber with a new batch of primitive magnesian magma. thus unit 512 represents the residing magma which in principle grades continuously into the mixed magma at the evolved end of the mixing curve. this also explains the local co-occurrence of lavas of unit 511 and 512, as in the profiles in kuugannguaq (fig. 9), asuutaa and killerpaat qaqqarsuat (fig. 10). petrography. the most magnesian rocks of the skarvefjeld unit (with mgo >10 wt%) are strongly olivine-phyric and similar to picrites of the vaigat formation. these lavas are confined to a small area around orlingasoq on 1800 1600 1400 1200 1000 800 600 400 200 1800 mm 1600 1400 1200 1000 800 600 400 200 0 km ls ls ls ls ls ls ls ls ls ls ls ls ls ls ls ls ls ls ls in ls ls ls ls lst? t? t? ls ls ls ls ls fe fe fe fe fe 1 2 nunngarut coal mine peak 1727 m killerpaat qaqqarsuat killerpaat qaqqarsuat inussuk q u l l i s s a t 3 4 5 6 7 8 9 10 11 km gv03_03_033_04_lml.eps se nw in invasive lava fe native iron ls landslipped area houses (qullissat) m ali gâ t f or m at io n va iga t f or m at io n at an ike rlu k fm t? ordlingassoq member paleocene? cretaceous fluvial sandstone, age uncertain atane formation, sandstone with shale beds and coal layers paleocenequaternary undifferentiated maligât formation subaerial lava flows olivine-phyric picrites (unit 450)subaerial lava flows foreset-bedded hyaloclastites intrusion of iron-bearing magnesian andesite, probably equivalent to asuk member undifferentiated landslipped volcanic rocks assoq member sandstone and mudstone nâujat member mudstone foreset-bedded hyaloclastites lower rinks dal mb lava flows with prominent entablatures and thick top breccias, partly or wholly subaqueous (unit 509) akuarut unit lava flows. iron-titanium-rich basalt (unit 513) lower rinks dal mb subaerial lava flows. feldspar-phyric and aphyric basalt (unit 512) boundary between massive lower part and thick top breccia of lava flow rinks dal member skarvefjeld unit (unit 511) olivineand feldspar-microphyric basalt ice and perennial snow undifferentiated deposits a n a n n n 509 511 512 513 514–517 450 fig. 48. photogrammetrically measured section along the north coast of disko around qullissat. a pencil drawing by r.r.j. hammer of the same coastal section (from steenstrup 1900) is shown for comparison. numbers are lithological unit codes. 67 north-eastern disko where they are intercalated with less magnesian lavas of the same unit (fig. 10, profiles 1, 2). basalts with 7–10 wt% mgo are characterised by a few equidimensional olivine phenocrysts and abundant skeletal, frequently platy, olivine microphenocrysts, and by microphenocrysts of plagioclase which increase in abundance as mgo decreases to 7 wt% (fig. 50a). tiny semitransparent chromite crystals occur enclosed in olivines in these rocks. basalts with 6–7 wt% mgo have augite as an additional phenocryst phase to olivine and plagioclase, and some rocks contain glomerocrysts of augite, plagioclase and olivine (fig. 50b). 1800 1600 1400 1200 1000 800 600 400 200 1800 mm 1600 1400 1200 1000 800 600 400 200 0 km ls ls ls ls ls ls ls ls ls ls ls ls ls ls ls ls ls ls ls in ls ls ls ls lst? t? t? ls ls ls ls ls fe fe fe fe fe 1 2 nunngarut coal mine peak 1727 m killerpaat qaqqarsuat killerpaat qaqqarsuat inussuk q u l l i s s a t 3 4 5 6 7 8 9 10 11 km gv03_03_033_04_lml.eps se nw in invasive lava fe native iron ls landslipped area houses (qullissat) m ali gâ t f or m at io n va iga t f or m at io n at an ike rlu k fm t? ordlingassoq member paleocene? cretaceous fluvial sandstone, age uncertain atane formation, sandstone with shale beds and coal layers paleocenequaternary undifferentiated maligât formation subaerial lava flows olivine-phyric picrites (unit 450)subaerial lava flows foreset-bedded hyaloclastites intrusion of iron-bearing magnesian andesite, probably equivalent to asuk member undifferentiated landslipped volcanic rocks assoq member sandstone and mudstone nâujat member mudstone foreset-bedded hyaloclastites lower rinks dal mb lava flows with prominent entablatures and thick top breccias, partly or wholly subaqueous (unit 509) akuarut unit lava flows. iron-titanium-rich basalt (unit 513) lower rinks dal mb subaerial lava flows. feldspar-phyric and aphyric basalt (unit 512) boundary between massive lower part and thick top breccia of lava flow rinks dal member skarvefjeld unit (unit 511) olivineand feldspar-microphyric basalt ice and perennial snow undifferentiated deposits a n a n n n 509 511 512 513 514–517 450 distribution. the distribution of the skarvefjeld unit is shown in fig. 15. the skarvefjeld unit is centred on disko east of the gneiss ridge but is also present on western disko and south-eastern nuussuaq (fig. 15; larsen & pedersen 1990, fig. 4). on eastern and southern disko the unit is easily distinguished lithologically as a compact package of foreset-bedded hyaloclastites capped by their equivalent subaerial facies of thin pahoehoe lava flows. with this morphology the unit has been mapped on the 1:100 000 scale geological maps uiffaq 69v.1s and pingu 69v.2n (unit βfph2), as well as on the 1:20 000 geological sections (south disko section, central disko section). the unit is well exposed on the south coast of disko between skarvefjeld and ippik where it is up to 180 m thick, and on eastern disko around kvandalen and blåbærdalen where thicknesses are 50–150 m. it is mf42_17cmimg_2738evsmod_c.eps 450 509 511–512 513 fig. 49. facies of the maligât formation at the north-western shore zone of the assoq lake basin. subaerial picrite flows of the vaigat formation unit 450 are overlain by two voluminous lava flows of maligât formation unit 509 with very prominent dark grey entablature zones and light grey brecciated tops, indicating emplacement into water. these are followed by subaerial flows of units 511–513. the height of the exposed section is c. 600 m. killerpaat qaqqarsuat, north disko (north-western part of the coastal section in fig. 48). photo: erik vest sørensen. 5 mm mf43a_8cm362315polcut_c.eps a 1 mm ol pl glass cpx mf43b_8cm176765g_c.eps b fig. 50. thin sections of rocks from rinks dal member unit 511, skarvefjeld unit. a: typical basalt lava with a few equidimensional olivine phenocrysts and abundant skeletal, commonly very elongate olivine phenocrysts (all altered to yellowish brown clay) together with a few plagioclase glomerocrysts (white) in a medium-grained groundmass. scanned thin section. sample ggu 362315 (mgo = 7.3 wt%), laksedalen, eastern disko. b: pillow breccia of basalt with glomerocrystic plagioclase (pl), fresh olivine (ol) and augite (cpx) in a glassy groundmass with tiny crystals of the same phases. microphotograph, plane polarised light. sample ggu 176765 (mgo = 6.5 wt%), skarvefjeld, south disko. 68 69 also present on northern disko around qullissat where it is c. 200 m thick and shows a morphological variety comprising subaqueous lava flows, foreset-bedded hyaloclastites and associated thin subaerial pahoehoe lava flows, and thick subaerial lava flows (figs 48, 49). on central and northern disko, lavas of the skarvefjeld unit are in subaerial facies and lithologically indistinguishable from the surrounding lavas. two to five flows referred to the skarvefjeld unit are present in nearly all profiles from orpiit qaqqaat to daugaard-jensen dal, sorte hak, kuugannguaq, orlingasoq and qullissat, commonly intercalated with flows referred to unit 512, as explained above. on southern nuussuaq one to two water-influenced flows are found lowest in the profiles at giesecke monument and eqi. lithologies. the volcanic rocks of the skarvefjeld unit occur as subaerial and subaqueous lava flows, pillow lavas, hyaloclastites and volcaniclastic beds. the successions of thin pahoehoe flows and foreset-bedded hyaloclastites, to which all the picritic rocks belong, show considerable lithological similarity to the picritic rocks of the vaigat formation. many lavas form flow fields consisting of many thin flow lobes with ropy flow surfaces, testifying to the fluid character of the magma (fig. 51). subaerial lava flows with mgo >10 wt% tend to develop subhorizontal, vesiculated segregation veins (fig. 52). soil layers of centimetre to decimetre thickness within the succession show that the skarvefjeld unit was erupted in several phases separated by sufficient time intervals to allow soil formation. mf44_8cmakp_1987_10_24_c.epsfig. 51. well-preserved ropy surface of a thin pahoehoe lava flow in the skarvefjeld unit. south-western slope of qinngusaq mountain, innermost head of kvandalen valley, east disko. mf45_8cmakp_1987_09_19modcut_c.eps fig. 52. a 7 m thick subaerial pahoehoe flow of olivine-rich basalt of the skarvefjeld unit (511) with subhorizontal, vesiculated segregation veins. the overlying flow is a feldspar-phyric basalt also of unit 511. ridge between frederik lange dal and charles polaris dal, east disko (fig. 10, profile 9 around 600 m). the background shows the same succession in the north wall of charles polaris dal. for location, see fig. 140. mf46_8cmakp_1987_09_26cut_c.epsfig. 53. pillow lava of the skarvefjeld unit: basalt pillows in a brownweathering matrix of volcanic glass, formed close to the entry point of the flow into the assoq lake. length of hammer 32 cm. north side of frederik lange dal, east disko; for location, see fig. 140. 7070 where the pahoehoe flows entered the assoq lake, pillow lavas (fig. 53) and hyaloclastites were formed, and volcaniclastic sediments occur as distal deposits along the basin floor (fig. 54). these sediments frequently contain scattered fragments of silicified wood or coal. mudstone of the assoq member is occasionally exposed beneath the volcanic rocks (fig. 55). the skarvefjeld unit formed the western shore zone of the assoq lake for nearly 100 km from south to north across disko (fig. 15). it prograded into the lake in a general easterly direction until the magma was exhausted. despite the fact that much of unit 511 is composed of pahoehoe flow fields which are unlikely to have flowed individually for more than at most a few tens of kilometres and hence must have been erupted on the present disko, no trace of any feeder dykes or necks have been located. three subareas of inflow into the assoq lake can be recognised and are described below. south coast of disko around skarvefjeld the south coast of disko around skarvefjeld has been described by heinesen (1987), larsen & pedersen (1990) and documented by pedersen et al. (2000) and in the south disko section. the skarvefjeld unit is exposed in the coastal cliffs over a distance of c. 7 km from skarvefjeld in the west to the western side of brededal in the east (south disko section at 84.3–91.3 km). here it forms a 80–100 m thick succession of foreset-bedded hyaloclastites and an overlying 30–40 m thick succession of pahoehoe flow lobes with transitional breccia facies at the interface between lavas and hyaloclastites (fig. 7, mf47_8cmakp_1987_10_19cut2_c.eps mf48_8cmakp_1987_10_14cut2_c.eps fig. 54. relatively fine-grained hyaloclastite with a few larger clasts deposited as a sediment gravity flow distally from the entry point of the lava flow into the assoq lake. about 15 cm of the underlying assoq member siltstone are seen lowest in the photo. length of hammer 55 cm. innermost head of kvandalen, east disko; for location, see fig. 140. fig. 55. laminated, silt-streaked mudstone with normal and inverse grading (shown by white open triangles at left) deposited from distal density currents, overlain by nearly homogeneous mudstone (above dashed line). assoq member underlying the skarvefjeld unit, detail of loose boulder; diameter of lens cap 5 cm. innermost head of kvandalen, east disko; for location, see fig. 140. 71 profile 4). the foresets dip consistently s to se, showing that the subaerial lavas flowed into the assoq lake basin from the north and north-west. individual foresets can be followed from the palaeoshore to the basin floor and demonstrate a lake depth of 80–100 m at the time (figs 56–57, 89). the hyaloclastites are separated from the underlying subaqueous lava flows of unit 509 by an up to a few decimetres thick veneer of dark mudstone, which may be disturbed by the hyaloclastites. the western and eastern margins of unit 511 are unexposed due to scree or erosion. at ippik the top of the subaerial lava flows of the skarvefjeld unit is eroded and covered by a c. 2 m thick bed of fine-grained yellow sandstone with millimetresized clasts of quartz and chert but entirely devoid of tertiary volcanic components. the sandstone is a distal facies of sandstone from the assoq member (dam et al. 2009) which is well exposed at tuapaat qaqqaat some 25 km to the east (fig. 41; also fig. 72). the sandstone at ippik is covered by lava flows of unit 512 (figs 57–58). all the basalts of the skarvefjeld unit along the south coast of disko show a narrow compositional range (6.1–6.5 wt% mgo) and could have originated from the same flow field. east disko around kvandalen the lavas and hyaloclastites of the skarvefjeld unit are shown as unit βfph2 on the map sheet pingu (pedersen et al. 2001). on the central disko section at 79.5–93 km, they comprise units m5 (foreset-bedded hyaloclastites) and m6 (lava flows) in the inner part of kvandalen. three localities aligned nw–se for about 30 km (inner kvandalen, charles polaris dal and point 1123 m) illustrate the compositional and lithological variation of the skarvefjeld unit on eastern disko. mf49_17cmakp_1987_01_29mod3_c.eps 514–518 511 hy 513 509 sla fig. 56. systematic upward changes of volcanic facies in the succession beneath skarvefjeld, demonstrating the filling of the assoq lake basin. the lowest unit (509) is 160 m thick and comprises at least four subaqueous lava flows with thick, yellowish brown breccia tops (509 sla); deposits of mudstone have accumulated in the lows on the undulating surface of the lowest flow (arrow). the overlying skarvefjeld unit comprises an 80 m thick horizon of foreset-bedded hyaloclastite (511 hy; the foresets dip towards the viewer); the associated overlying subaerial pahoehoe flows are concealed beneath scree. scree also conceals a few flows of unit 512 and a thin horizon of quartzo-feldspathic sandstone, probably a fluvial deposit. the following units 513 to 518 are all in subaerial facies with red-oxidised top rubble. cliff face, 920 m high, beneath skarvefjeld, south coast of disko just east of qeqertarsuaq town; see also figs 57, 58, 89. 7272 ippik 400 m sst 297 m pahoehoe flows (511) sst sst irregular flows (513) subaerial flows oxidised top (513) thick flow with colonnade+entablature (512) 300 m foresetbedded hyaloclastite breccia (511) three subaqueous lava flows(509) mudstone schlieren 200 m mf50_ippikscan_bitmap.eps fig. 57. photogrammetric interpretation of the volcanic succession at ippik 4 km east of skarvefjeld, south coast of disko. subaqueous lava flows of unit 509 and hyaloclastites of unit 511 filled in the assoq lake. the foresets can be traced from top to bottom of the hyaloclastite, indicating a lake depth of around 85 m. note two very thin levels of sandstone (sst). lithological unit codes in parentheses. see also fig. 58. from heinesen (1987), with slightly modified text. mf51_17cmakp_1989_02_11mod_c.eps 512 la sst sst 511 hy 511 la 513 513 fig. 58. the volcanic succession at ippik. hyaloclastites of the skarvefjeld unit (511 hy) and subaerial lava flows of units 511, 512 and 513. the single flow of unit 512 has a well-developed colonnade indicating emplacement on a wet surface; thin layers of quartzo-feldspathic sandstone (sst), probably fluvial, are seen both below and above this flow. see also fig. 57. ippik, south coast of disko (south disko section at around 90 km). 73 inner kvandalen. fig. 10, profile 8. in the innermost part of kvandalen the skarvefjeld unit is about 200 m thick. it overlies more than 20 m of assoq member mudstones, which contain plant fossils and thin shell fragments of a small bivalve. the lowermost c. 180 m of the unit is in subaqueous facies and forms various hyaloclastites and volcaniclastic beds, which were derived from at least three different eruptive events separated in time by beds of non-volcanic clastic sediments. the lower hyaloclastite unit comprises some of the most evolved basalts of unit 511 (with c. 5.5 wt% mgo) and is composed of glassy clasts and pillow fragments of plagioclase-augite microphyric basalt together with clasts of subaerial basalt, mudstone and scarce siderite concretions derived from clastic sediments. the unit comprises several beds with varying amounts of moderately baked mudstone and conglomeratic layers of up to decimetresized basalt clasts derived by coastal erosion of subaerial lava flows of unit 511. some of the beds are dominated by volcaniclastic gravel and are separated by up to decimetre-thick beds of deformed mudstone (fig. 54). a middle unit is composed of slightly more magnesian hyaloclastite (with c. 6 wt% mgo) with scattered clasts of mudstone. the uppermost unit is up to 100 m thick and consists of foreset-bedded hyaloclastite and pillow lava of distinctly more magnesian basalt (with 8–9 wt% mgo) which were fed from overlying, compositionally similar subaerial pahoehoe lava flows. charles polaris dal. fig. 10, profile 9. about 11–12 km south-east of profile 8, the skarvefjeld unit is excellently exposed on both walls of charles polaris dal, in particular the south wall at the corner to frederik lange dal. the skarvefjeld unit is here about 125 m thick and overlies quartzo-feldspathic sandstone and c. 80 cm of poorly exposed mudstone of the assoq member. the lowermost part of unit 511 is an about 7 m thick volcaniclastic bed composed of glassy basalt clasts and pillow fragments together with scattered millimetre-sized polished quartz grains, rare centimetre-sized clasts of polished chert, and mf52_17cmlml_1987_03_27_c.eps 511 hy 511 la 513 512 sstsst fig. 59. the skarvefjeld unit (511) in the south wall of charles polaris dal, eastern disko. thick foreset-bedded hyaloclastites (511 hy) showing infill from the west (right), and associated subaerial thin pahoehoe flows ending with one or two thicker flows (511 la). a c. 10 m horizon of quartzo-feldspathic sandstone (sst) separates the overlying flows of unit 512 and 513. the subaerial part of the skarvefjeld unit (511 la) is c. 60 m thick. for location, see fig. 140. 7474 fragments of coal and silicified wood picked up from the underlying sediments. this bed is overlain by a 7 m thick columnar-jointed basalt of unit 511 with poorly exposed lower and upper contacts, interpreted as an invasive flow. this is overlain by a c. 45 m thick, foreset-bedded hyaloclastite showing inflow from the west, covered by a c. 50 m thick succession of thin pahoehoe lava flows feeding the hyaloclastites (fig. 59). the upper part of the skarvefjeld unit in charles polaris dal is composed of two olivine-rich lava flows with up to 12.2 wt% mgo and a top flow of feldspar-phyric basalt (fig. 52). the olivine-rich flows have very dark grey, partly olivine-cumulative basal zones with prominent sub-horizontal segregation veins, and upper light grey zones, which are vesicular and less rich in olivine. the top zones of both flows are high-temperature oxidised and the upper flow has a c. 5 cm cover of lateritic soil. the upper picritic flow shows a distinct local up-doming and a thickening of its olivine-rich lower part which might indicate proximity to an eruption site, but no trace of a volcanic feeder has been recognised. the uppermost flow of unit 511 is a c. 7 m thick feldspar-phyric basalt with red-oxidised top with traces of lateritic soil. it is covered by a thick lava flow of unit 512. south wall of kvandalen east of point 1123 m. fig. 7, profile 12. the easternmost exposures of the skarvefjeld unit are situated on the south wall of kvandalen about 29 km south-east of profile 8. at least 55 m of the skarvefjeld unit are exposed in a slightly slipped block c. 1.5–2 km east of point 1123 m (fig. 60; larsen & pedersen 1990, fig 7; dam et al. 2009, fig. 131). the lower boundary is mf53_17cmakp_1987_11_24_c.eps 511 hy 514 515 516 sst sst 511 la 512 la 513 inv fig. 60. the easternmost exposures of the skarvefjeld unit in kvandalen. the view is from west-south-west (point 1123 m) and shows hyaloclastites (511 hy) and subaerial lava flows (511 la) of the skarvefjeld unit, in total c. 55 m thick. units 512–516 are each represented by a single flow. the flow of unit 513 has yellow quartzo-feldspathic sandstone (sst) at its base and top and is probably invasive (inv). see also dam et al. (2009, fig. 131). the exposure is slightly slipped. the white outcrops in the background are cretaceous sandstones in kvandalen. south side of kvandalen c. 1.5 km east of point 1123 m, east disko (fig. 7, profile 12). 75 not exposed. the skarvefjeld unit here comprises 20–25 m of foreset-bedded hyaloclastites covered by c. 30 m of subaerial pahoehoe flows feeding the hyaloclastites. the compositional variation is much smaller than in the localities described above (8.0–8.2 wt% mgo in the hyaloclastites and the lavas just above them, and about 6.5 wt% mgo in the uppermost lava flows), which suggests that the skarvefjeld unit here is distal relative to its source and represents just one or two flow fields. the upper boundary of the pahoehoe lavas is a planar erosion surface covered by about 10 cm of coaly mudstone. the final volcanic front, which marks the exhaustion of unit 511 is not exposed at any locality. north coast of disko around qullissat the northern part of the assoq lake deposits and volcanic infill are virtually unexposed over a distance of c. 30 km along the vaigat coast between innermost kvandalen and qullissat. the lack of exposures is caused by large landslides, which have covered both the lowermost part of the maligât formation and the underlying cretaceous to tertiary sediments (pedersen et al. 2001; central disko section). however, blocks of hyaloclastites of the skarvefjeld unit occur in the landslipped masses at several localities, and the unit may therefore be present along the whole distance. exposures of the skarvefjeld unit reappear in the coastal mountain walls at inussuk and qullissat (fig. 10, profiles 4, 5; fig. 48). over a distance of about 8 km the near-vertical walls of the inussuk mountain in the south-east and the killerpaat qaqqarsuat mountain in the north-west show an up to 200 m thick succession of subaerial and subaqueous lava flows and foreset-bedded hyaloclastites of the skarvefjeld unit. the basalts have a limited chemical variation (6–7 wt% mgo) but form both thick individual flows and successions of thin pahoehoe tongues. the repeated facies variations, with two hyaloclastite horizons (fig. 48 at 2.5–5 km), demonstrate that the basalts were emplaced in a number of individual eruption events that were sufficiently spaced in time to record synvolcanic subsidence. in the killerpaat qaqqarsuat mountain, only the lowermost lava flow of the skarvefjeld unit records wet conditions at the time of inflow. this flow overlies partly subaqueous flows of unit 509 (see section on unit 509) and has an intensely jointed colonnade and entablature but a vesiculated and oxidised top surface and a thin coat of lateritic soil. a few kilometres to the south-east mf54_17cmimg_2681evsmodcut_c.eps 511 la la hy 512 511 hy 509 sla 511 sla 513 fig. 61. the eastern part of the mountain wall inussuk just south of qullissat, north-eastern disko (fig. 48 at 3–4.2 km). the major part of the wall comprises the skarvefjeld unit in subaqueous lava facies (511 sla), foreset-bedded hyaloclastite facies (511 hy) and subaerial lava facies (511 la). the foresets of the lower hyaloclastite horizon are nw-inclined whereas the foresets of the upper hyaloclastite horizon are se-inclined (fig. 48). the total thickness of unit 511 is here c. 200 m. photo: erik vest sørensen. 7676 the same flow is probably present as one or two entirely subaqueous, thick lava flows emplaced on top of subaqueous flows of unit 509 (figs 10, 61). the palaeo-water depth cannot be estimated from these flows, but they are covered by a foreset-bedded hyaloclastite horizon fed from a flow field of pahohehoe lava lobes, and the 35–45 m high foresets truly reflect the water depth at that time. after the filling of the assoq lake, the area was inundated again and filled by a new series of pahoehoe flows of the skarvefjeld unit, which form an upper, just 5–10 m thick, foreset-bedded hyaloclastite. this marks the final filling of the assoq lake in this part of disko, and the final flows of unit 511 are entirely subaerial (fig. 61). unit 512 (lower transition flows) composition and petrography. these flows form a transition from the skarvefjeld unit to the overlying fe-ti-rich akuarut unit. in a plot of tio2 vs mg-number (see fig. 97) they form a partially more evolved extension of the trend of the skarvefjeld unit, with equal or higher tio2 contents (2.4–3.2 wt%) and lower mg-numbers caused by higher feo* and relatively low mgo (5.8–6.8 wt%). a few exceptions with lower tio2 are found on eastern disko. the lower transition flows are moderately phyric to almost aphyric basalts with some phenocrysts, commonly glomerocrysts, of plagioclase up to 6 mm in size, scarce augite up to 1 mm in size, and olivine up to 1 mm in size but always pseudomorphed by smectite (fig. 62). chromite has not been observed. distribution. the lower transition flows are found throughout disko; on eastern disko the unit oversteps unit 509 and the skarvefjeld unit (511) and forms invasive flows which are the oldest in the area (fig. 7, profiles 8, 10, 11, 13, 14). on most of disko unit 512 is 20–100 m thick and consists of only one to three flows; an exception is seen at qinngusaq on north-eastern disko where a succession of four flows reaches 170 m thickness and two flows have a combined thickness of 100 m because of ponding (fig. 10, profile 8). unit 512 is present in the southernmost part of nuussuaq where it is represented by a single invasive flow, the lowest one in the profiles at tartunaq, umiusat and eqip qaqqaa, and flow 2 at giesecke monument (fig. 11). transition flows are absent in a few profiles (lyngmarksfjeld, orlingasoq). lithology. the lower transition flows are widespread, massive, blocky pahoehoe lava flows, and except for their chemical composition, they are indistinguishable from most other rinks dal member basalts (fig. 63). at most localities, the one to four flows of unit 512 are situated between the skarvefjeld unit (511) and the overlying prominent akuarut unit (513). however, at kuugannguaq se and killarpaat qaqqarsuat (fig. 9, profile 5; fig. 10, profile 4) unit 512 lavas interdigitate with flows of both units 511 and 513. on eastern disko (frederik lange dal, blåbærdalen and point 1123 m in kvandalen) unit 512 is represented by a single 20–30 m thick subaerial lava flow, which overlies subaerial pahoehoe lava tongues of the skarvefjeld unit (fig. 60). the flow top is eroded and covered by a widespread, around 10 m thick horizon of non-volcanic, quartzo-feldspathic sand and siltstone of the assoq member (fig. 60; dam et al. 2009, fig. 131). at frederik lange dal this horizon contains marine dinoflagellate cysts in some samples (piasecki et al. 1992), indicating marine incursions into the assoq lake at this stage. the mf55_8cm340873polcut_c.eps cpx pl 5 mm ol fig. 62. thin section (scanned) of a typical lava from rinks dal member unit 512, lower transition flows. basalt with many glomerocrysts and individual phenocrysts of plagioclase (pl) and less common augite (cpx) and olivine (ol, altered to dark brown clay), in a finegrained matrix. sample ggu 340873, qinngusaq, north-east disko. 77 sediment horizon on top of unit 512 extends for more than 50 km to the south coast of disko and farther westwards where it thins gradually; a sand layer a few decimetres thick persists into the lava plateau at skarvefjeld and ippik (fig. 7, profile 4; fig. 57). invasive lava flows on eastern disko the subaerial lava flows of unit 512 are nowhere directly observed to pass onto or into the assoq member sediments, but basaltic sill-like bodies in the sediments are widespread on eastern disko. they have chemical compositions similar to unit 512 lavas and occur at an equivalent stratigraphic level, and they are interpreted as invasive lava flows (larsen & pedersen 1990). they occur from tuapaat qaqqaat on the south coast to pingu and inngigissoq on the east coast (fig. 7, profiles 8–14; fig. 10, profiles 10–12; fig. 64). at inngigissoq the upper of two such lavas is thoroughly pillowed (fig. 7, profile 14), indicating that the sediments were wet and unconsolidated at the time of intrusion. invasive lava flows on southern nuussuaq a very well-exposed invasive lava flow of unit 512 (figs 65, 66) can be followed semi-continuously for about 18 km along the south coast of nuussuaq from giesecke monument in the west to tartunaq in the east (south nuussuaq section at 59–77 km). samples from three profiles (fig. 11, profiles 7, 9, 10) have virtually identical chemical compositions, confirming that the flow represents a single eruption. the flow burrowed into unconsolidated mud of the assoq member (formerly named the aussivik member). figure 66 shows that at c. 72.7 km the boundary between sandstone of the umiussat member and mudstone of the assoq member is located below unit 512. farther to the east unit 512 is invasive into the umiussat mb (fig. 11, profile 10). unit 512 is at least locally composed of two columnar-jointed bodies mf56_17cmlml_1991_14_06mod4_c.eps 507 hy 507 la 509 511 512 513 513 515 514 fig. 63. the eastern part of the type section for the rinks dal member at orpiit qaqqaat. the photo corresponds approximately to fig. 24 at 3–4 km but is viewed at a different angle. from unit 509 upwards, all flows are subaerial. the sample profile follows the left edge of the gully (fig. 24). the fe-ti-rich akuarut unit (513) is recognised by having stronger reddish-brown colours than the other units. see also figs 25 and 31. south side of the orpiit qaqqaat mountain, south-central disko. 7878 with upper and lower contacts chilled against mudstone (pedersen 1975b; see also fig. 76). the flow is 20–50 m thick along the coastal section and extends for more than 5 km inland (pedersen et al. 2007b); it has an estimated volume in excess of 2 km3. this is a much larger volume than that of the preceding thin lavas of unit 511, and the flow evidently had volume enough to extend the lava area much farther into the assoq lake basin. a separate sill-like body which covers at least 4 km2 c. 5 km north-north-east of giesecke monument (pedersen et al. 2007b), represents the northernmost extension of unit 512. mf57_17cmimg_2427evsmodcut_c.eps 514 inv 512 inv a u p 515 fig. 64. sediments and lava flows on eastern disko. two invasive lava flows belong to units 512 and 514, whereas unit 513 is absent. the flows of unit 515 are in normal subaerial facies (fig. 10, profile 12). the sediments belong to the atanikerluk formation. the boundary between the upper part of the pingu member (p) and the umiussat member (u) is indicated. the lower boundary of the assoq member (a) is not well exposed in this section, but is probably located above 512 inv. the north-east facing slope of inngigissoq, east disko. the outcrop is shown on the central disko section around 103 km (viewed at a different angle) and in dam et al. 2009, fig. 132. photo: erik vest sørensen. 79 fig. 65. invasive lava flows of units 512 and 513 in mudstones (the 512 flow) and sandstones (the 513 flow) of the assoq member. flows of units 514–515 are in normal subaerial facies. the sandstone seen in the lower part of the photo is the umiussat member. the fine-grained sandstones are interpreted as low-energy fluvial deposits. slope beneath point 975 m, south coast of nuussuaq (fig. 66 at c. 71.8–72.5 km). fig. 66. photogrammetrically measured section of the south coast of nuussuaq around point 975 m, showing invasive lava flows of units 512 and 513 and subaerial lava flows of unit 515(?). blue three-digit numbers are lithological codes. the invaded sediments are sandstones (yellow, labelled tsa) and mudstones (dark grey) of the assoq member of the atanikerluk formation. lower sediments are mudstones of the naujât member (tmu) of the atanikerluk formation and sandstones of the kingittoq member (k) of the cenomanian atane formation. ls: landslipped blocks. δ: dykes. the dark purple area at atanikerluk is an eocene sill. excerpt from the south nuussuaq section (pedersen et al. 1993). note that the horizontal scale is at the top. mf58_17cmakp_1986_04_01mod2_c.eps 514-515 la point 975 m 513 inv 512 inv 513 71 72 73 74 km 200 400 600 800 1000 1200 sm sm k tsa tmu mt ml point 1033 m atanikerluk point 970 mpoint 975 mkeglen assoq mb umiussat mb naujât mb kingittoq mb ls ls δ δ ls gv03_03_054_lml 515? 513 512512 nw se not exposed 8080 middle rinks dal member summary of the main features of the middle rinks dal member • comprises a single unit, the akuarut unit (513), with high tio2 (≥3.2 wt%). • stratigraphic ‘backbone’ of the rinks dal member. generally mappable by a combination of chemistry and strong brown colours. not well distinguished on north-west disko and not shown on the 1:100  000 map sheet qullissat. otherwise shown on all relevant geological maps and photogrammetric sections. • covered the last remaining hills of the disko gneiss ridge on southern disko. • spread on a flat plain on disko and stepped farther north onto the vaigat formation picrite shield on south nuussuaq. • all lava flows are subaerial. on eastern disko and southern nuussuaq they are also invasive into assoq member sediments. rare dinoflagellate cysts in the assoq member indicate episodic marine incursions on disko. • the most evolved basalts of the rinks dal member. tilted ree spectra suggest garnet fractionation in deep-seated magma chambers. unit 513 (akuarut unit) the akuarut unit is an important stratigraphic marker and is shown as unit βi on the three 1:100  000 scale geological map sheets covering disko south of 70°n (mellemfjord, uiffaq and pingu) and as unit mi on the 1:100 000 geological map sheet paatuut covering southern nuussuaq. it is also shown on the 1:20 000 scale geological sections showing the south coast of disko, central disko and the south coast of nuussuaq. composition. the flows of this unit are characterised by high contents of iron and tio2, and the unit was called the ‘fe-ti unit’ by larsen & pedersen (1990) and ‘irontitanium-rich basalts’ on the maps and sections. it is here defined as all flows around the middle level of the rinks dal member with tio2 ≥3.2 wt% (table 2; fig. 21). this value was chosen because it is practically applicable and results in the clearest divisions between units. thus defined, the flows of the akuarut unit have 3.2–4.8 wt% tio2, 4.4–6.8 wt% mgo, and 12.3–16.8 wt% feo*. these flows are the most evolved basalts in the rinks dal member, yet their mg-numbers show a large overlap with other units (fig. 97). petrography. the basalts of the akuarut unit range from almost aphyric to distinctly porphyritic or glomeroporphyritic with phenocrysts of plagioclase, augite and olivine (always altered; fig. 67). almost aphyric basalts occur through the entire compositional range in tio2. cognate dolerite inclusions occur occasionally. some of the invasive lava flows have a doleritic groundmass texture. distribution. the distribution of the akuarut unit is shown in fig. 15. the akuarut unit is present over nearly all of disko but is missing in a small area around pingu on eastern disko (fig. 7, profile 14; pedersen et al. 2001; central disko section). in some places, e.g. at lyngmarksmf59_8cm279038dglcut_c.eps pl 5 mm cpx ol fig. 67. thin section (scanned) of a typical lava from rinks dal member unit 513, akuarut unit. basalt with a few small plagioclase phenocrysts and a glomerocryst of plagioclase (pl), augite (cpx) and pseudomorphosed olivine (ol) in a fine-grained matrix with a dark, glass-rich streak at top of photo. sample ggu 279038, killerpaat qaqqarsuat, northern disk 81 fjeld, the unit can be distinguished lithologically because the flows are very brown and have thick, purple top breccias, but commonly it is not clearly visually distinguishable. the unit is thickest on western and central disko where there are from five to about 20 flows present (fig. 68) with a combined thickness of 150–400 m. it is probable that the unit is up to 600 m thick on north-western disko, but due to lack of flow-by-flow sampled profiles in that area this cannot be ascertained. it is represented by a single lava flow at the mouth of giesecke dal on north-western disko (fig. 13, profile 6), whereas its distribution among the mountain tops on northern disko on both sides of the kuugannguaq valley is unknown due to lack of sampling. on eastern disko, the akuarut unit thins eastwards; it goes into invasive facies, is reduced to five to one flows and peters out. a single flow reached sortebærdalen, kvandalen and aqajaruata qaqqaa, but the unit never reached pingu. on south and south-east nuussuaq, the akuarut unit is present south of the central aaffarsuaq valley (fig. 15) where it comprises one to four flows (fig. 11). the flows onlap the surface of the vaigat formation, which rises towards the north and west. it is thickest at giesecke monument and eqip qaqqaa where it is represented by four thick flows of which the lowest two are invasive. one subaerial flow reached point 1722 m just west of ataata kuua as the lowest flow of the rinks dal member, resting directly on the flows of the vaigat formation, but none reached point 1888 m only 8 km northeast of point 1722 m. in the south-eastern profiles of umiusat and tartunaq a single thin invasive flow is present, the same in the two profiles. over much of disko the akuarut unit forms a single, well-defined unit. this is the case on southern and eastern disko as seen in the south disko section, e.g. fig. 36. however, in some profiles on central disko there are a few flows with less than 3.2 wt% tio2 intercalated in the akuarut unit (fig. 9, profiles 3, 4, 5). on nuussuaq, a group of flows with less than 3.2 wt% tio2 is intercalated within the upper part of the akuarut unit, as seen on the south nuussuaq section at 50–56 km and fig. 11. the intercalated lavas are considered to belong to unit 514 (the upper transition unit) described below. lithology. the akuarut unit is mainly composed of subaerial lava flows. on eastern disko and nuussuaq the flows are water-influenced and commonly invasive. the flows are mostly blocky pahoehoe flows, but aa flow morphology is also encountered. typical flow morphologies are seen in the steep wall above ippik (fig. 58). in parts of southern, central and eastern disko and nuussuaq the flows have a distinctive, rusty brown weathering colour (figs 20, 43), but in other areas, such as north-western disko, the flows are greyish brown and do not show sufficient colour contrast to the surrounding flows to allow visual identification of the akuarut unit. individual lava flows extend over large areas and individual flow volumes must commonly exceed 1 km3. in general, the mean flow thickness is about 20 m on western disko and more than 30 m on central and eastern 16 17 18 19 km gv03_03_050_lml m4 i3 m8.6 m8 ? δmδm δm δm m8 m7.5 m7.4 perlertut kuuat 509-512 513 514-517 ? ?? ? wnw ese 1200 m 1000 800 600 400 200 fig. 68. the rinks dal member at perlertut, north coast of nordfjord, west disko, characterised by a large number of relatively thin lava flows. white and blue three-digit numbers are lithological codes; other annotations are from the original. some individual marker flows are shown with deviating colours. picritic lava flows of the vaigat formation are denoted i3. excerpt from the central disko section (pedersen et al. 2005). 8282 disko. some individual flows can be traced photogrammetrically on southern and central disko, and the sectioned length of these is 3.7–12.5 km (south disko section, central disko section). if a typical flow occupies a circular area with a diameter of 10 km, this amounts to c. 80 km2, and with a thickness of 20 m, it has a volume of c. 1.6 km3. one or two flows can be distinguished by a particular chemical composition; of these, one flow on nuussuaq extends over 17 × 10 km between six profiles with an estimated volume around 3.4 km3. another flow on south-western disko appears to be present over 55 × 35 km in five profiles; if it is truly a single flow its volume is in the order of 30 km3. no primary eruption sites or dykes have been found and it is assumed that the akuarut unit was erupted on western disko or in the western offshore areas, but eruption sites on central and south-central disko cannot be excluded. the time intervals between eruptions of the flows of the akuarut unit were frequently large enough to allow deep weathering, erosion of the lava tops and deposition of decimetre-thick layers of lateritic soil, well exemplified by lava flows in the killiit area (fig. 7, profile 2; fig. 69). the akuarut unit seems to have spread over a flat plain very close to sea level on disko and to have onlapped the low shield of vaigat formation picrites on nuussuaq. to the east, it was bordered by the shallow assoq lake basin. the lava flows progressed from a westerly direction and invaded and interacted with the sediments of the basin over a more than 100 km long shore zone extending from the south coast of disko to the aaffarsuaq valley on nuussuaq (fig. 16). on disko, dinoflagellate cysts have been recovered from interbasaltic sediments associated with the akuarut unit at two localities, indicating a brackish environment, whereas similar sediments on nuussuaq appear to be non-marine (piasecki et al. 1992). mf62_17cmimg_4493_c.eps laterite gneiss 513 513 fig. 69. two lava flows of the akuarut unit (513) onlapping the disko gneiss ridge. the lower flow was weathered and eroded and an up to 1 m thick brick-red laterite horizon was developed before the upper flow arrived. south coast of disko above killiit (south disko section at 68 km). 83 exposures of the akuarut unit in the eastern shore zone are scarce on much of eastern disko and south-eastern nuussuaq because of extensive landslides and solifluction, but the lithological variations are well-exposed in an 8 km long section along the south coast of disko between marraat qaqqaat and tuapaat qaqqaat, and also in a 24 km long section along the vaigat coast on nuussuaq between point 1760 m and tartunaq. these are described in the following. south coast of disko between marraat qaqqaat and tuapaat qaqqaat east of brededal good partial exposures of the contacts between lava flows and the largely contemporaneous sediments of the assoq member are seen over a distance of c. 8 km between marraat qaqqaat and tuapaat qaqqaat (south disko section at 108–116 km; fig. 41). here the akuarut unit interacted with the quartzo-feldspathic sands, which in this region form the upper part of the assoq member (dam et al. 2009). two exposures illustrate the interaction between subaerial lava flows and unconsolidated sand. rootless phreatic cones at marraat qaqqaat. at the slope of marraat qaqqaat (south disko section at c. 109 km; fig. 7, profile 7), the lowermost of three lava flows of unit 513 forms a c. 30 m thick invasive flow with a pronounced colonnade and a well-defined upper chill zone against a sandstone horizon which is no more than 5 m thick (fig. 70). above the sandstone, there is a 5–8 m thick mixed horizon consisting of a volcanic breccia with pillow-like clasts in a matrix of sand and glass fragments; the deposit may be called a peperite (fig. 71). this bed has a 20–30 cm thick, oxidised weathering surface. it is overlain by a 4–6 m thick, irregular, subaerial lava flow, which in places grades upwards into a number of spatter cones with beds of centimetreto decimetre-sized achnefig. 70. volcanic breccias interpreted as rootless cones. the lower lava flow (513 inv) is 30 m thick and has a columnar-jointed lower half with a sharp transition to entablature facies in its upper half. this flow has invaded a sandstone horizon (ss) only c. 5 m below the palaeosurface and re-erupted up through the sandstone, forming a volcanic breccia with sediment matrix (vbr 1). after a pause indicated by a red, weathering surface (red line), a 4–6 m thick lava flow arrived, and on top of this a brecciated zone of rootless phreatic cones (vbr 2) was created by explosions of heated vapour in the wet sand underlying the flow. see text for detailed descriptions. south coast of disko at marraat qaqqaat (south disko section at 109.2 km). mf63_17cma1984_02_06mod4cut.eps 513 inv vbr 1 vbr 2 ss ss 513 la 513 la 8484 liths (rounded pyroclastic fragments) at the top. this is covered by a massive, about 50 m thick subaerial lava flow with a several metres thick, blocky and scoriaceous top, which is the uppermost flow of the akuarut unit. this flow is in turn covered by 5 m of yellow sandstone. the breccia above the lower invasive lava flow is interpreted as basalt from the same flow, re-erupted through the very thin sandstone cover, while the spatter cones from the middle flow are interpreted as rootless phreatic cones created by explosions of heated vapour in the underlying wet sand. the oxidised weathering surface and the subaerial spatter cones show that the water was very shallow, probably a shore-near lake or a riverbed. thus, the assoq lake had shrunk and shallowed considerably since the formation of the rootless cones at niuluut, only 10 km east-south-east of marraat qaqqaat, beneath around 200 m of water in the lake at the time of deposition of unit 509 (larsen et al. 2006). invasive lava flows and conglomerate at tuapaat qaqqaat. about 6 to 7 km east of the rootless spatter cones at marraat qaqqaat, there are remarkable exposures of akuarut unit lava flows interacting with unconsolidated yellow sand of the upper part of the assoq member (fig. 41). as at marraat qaqqaat, there are three invasive flows. the two lower ones are compositionally identical to the two lower flows at marraat qaqqaat, but here they are separated by c. 50 m of sand (fig. 7, profile 8; fig. 72). the upper flow is compositionally identical to the 50 m thick upper flow at marraat qaqqaat, but here its thickness has decreased to 14 m and it overlies a thin mudstone horizon on top of 11 m of sand. the top of the flow is eroded and covered by several metres of a conglomerate with rounded, polished basalt clasts and a mixed matrix of quartzo-feldspathic yellow sand and basalt gravel (fig. 73). mf64_8cma1984_02_07modcut2_c.eps fig. 71. peperite deposit from rootless cones: a breccia consisting of volcanic clasts in a matrix of quartzo-feldspathic sand and glass fragments. length of hammer 32 cm. marraat qaqqaat, south disko. fig. 72. lava flows of the akuarut unit (513) invasive into sandstone of the upper assoq member. the fine-grained, friable sandstone shows cross-lamination indicative of deposition from low-energy currents, but the exposure is generally very poor. the sandstone between the two flows is c. 50 m thick, and more than 100 m of sand is indicated by patchy outcrops in the scree below the lower flow. ridge just west of tuapaat qaqqaat, south coast of disko (south disko section at 114.9–115.7 km, fig. 41). mf65_11cma890229mod_c.eps 513 inv 513 inv 513 inv 85 the presence of marine dinoflagellate cysts in the mudstone beneath the flow (piasecki et al. 1992) suggests that the conglomerate on top of the flow could have resulted from marine coastal erosion and abrasion. samples from the lower sand horizons at tuapaat qaqqaat do not contain dinoflagellate cysts. however, dinoflagellate cysts are also found in the mudstones at assoq that are invaded by flows of unit 509 (see text on unit 509). the dinoflagellate taurodinium granulatum, which is indicative of brackish-water conditions (fensome et al. 2016), is present at some levels (h. nøhr-hansen, personal communication 2017). the kinds and distribution of dinoflagellate cysts thus suggest that the assoq lake was subject to periodic marine inundations. the other known sediment locality with dinoflagellate cysts associated with subaerial lava flows of the akuarut unit is situated in frederik lange dal c. 40 km north of tuapaat qaqqaat. taken together and combined with the very extensive regional distribution of subaerial lava flows of the akuarut unit over most of disko and parts of southern nuussuaq (fig. 15), this supports the interpretation that the akuarut unit was erupted over a flat plain very close to sea level at that time. south coast of nuussuaq large ponded flow. the akuarut unit forms the base of the maligât formation between paatuut and point 1760 m west of giesecke monument (south nuussuaq section at 50–58 km) where it directly overlies picritic subaerial lavas and hyaloclastites of the vaigat formation. beneath point 1760 m (fig. 11, profile 5; fig. 74) the akuarut unit contains an up to 150 m thick, ponded lava flow that consists of a number of irregular, strongly columnarjointed basalt bodies with distinct but thin colonnades and very thick entablature zones. the flow has a top zone of highly vesiculated, blocky and scoriaceous basalt which clearly solidified subaerially. towards the east, the large flow overlies mudstones of the assoq member. the large flow is interpreted as having ponded in a lake basin, leading to temporary complete displacement of the water from the area. it is covered by a poorly exposed, 10–20 m thick subaerial lava flow also of unit 513. invasive lava flows along the vaigat coast. about 5 to 6 km south-east of point 1760 m, beneath the ridge just south of giesecke monument, more than 200 m of invasive lava flows of the maligât formation are exposed, interbedded with contemporaneous or slightly older mudstones of the lower part of the assoq member (fig. 11, profile 7, figs 74, 75). above two older invasive flows of units 511 and 512, there are two massive sill-like flows of the akuarut unit, which both have a chemical composition that differs significantly from that of the large ponded flow beneath point 1760 m. the lower flow is c. 60 m thick and was emplaced in a single pulse. it has a poorly exposed contact against assoq member mudstones at both the base and the top. the upper flow is 80–90 m thick and composed of at least two cooling units (note the columnar-jointing pattern in fig. 75); it has a bulging, irregular top composed of decimetre-sized pillow bodies. the flow has baked and hydrothermally affected the overlying mudstones, which include a c. 0.5 m thick coal-rich layer with a deposit of well-preserved, compacted, close-lying fossil leaves. the pillowed structure indicates that the flow invaded unconsolidated wet sediment. on top of the large invasive flow there is a lava flow of unit 514, which has a subaerial flow top; it is followed by two subaerial flows which form the upper part of the akuarut unit. along the vaigat coast south-east of giesecke monument the number of flows in the akuarut unit decreases, and south-east of umiusat there is only a single flow present. most contacts to sediments are covered by scree, but over a distance of 7 km from just west of keglen to the fig. 73. conglomerate overlying the thin top flow on tuapaat qaqqaat, south disko, consisting of polished, rounded basalt clasts in a matrix of basalt gravel and yellow, quartzo-feldspathic sand. marine dinoflagellate cysts in a thin mudstone beneath the lava flow suggest that the rounded cobbles and pebbles could have resulted from marine coastal erosion and abrasion. the ruler is 15 cm long. the conglomerate is indicated in fig. 41. mf66_8cma1984_10_05modcut2_c.eps 8686 56 55 54 57 58 59 60 k m 20 0 40 0 80 0 16 00 18 00 m 14 00 20 0 40 0 60 0 80 0 10 00 16 00 18 00 m 14 00 t s a sil l m tm l m n m m m l m tm m m l t m u in va siv e flo w e e ? ? i 3 i 3i 3i 3 i 1 i 2 po in t 1 58 0 m ip eq ar fiu nn gi at an e fm , q ila ki ts oq m b at an e fm , k in git to q m b at an ike rlu k fm , n au jât m b g ie se ck e m on um en t u pp all uk po in t 1 76 0 m ls ls ls ls in in k ls ls sil l sil l ls ls g v0 3_ 03 _0 53 _l m l 51 4 52 0 51 7 51 6 51 5 51 4 51 3 45 0 45 0 51 3 n w se n o t e x p o se d fi g. 7 5 fi g. 7 4. p ho to gr am m et ric al ly m ea su re d se ct io n al on g t he so ut h co as t o f n uu ss ua q at th e e as te rn m os t o cc ur re nc e o f t he v ai ga t f or m at io n (u ni t 4 50 ), sh ow in g h ya lo cl as tit es an d la va fl ow s a bo ve p oo rly ex po se d se di m en ts (m an y la nd sli ps , l s) . t he e as te rn m os t l av a o f t he v ai ga t f or m at io n is an in va siv e p ic rit e f lo w. f ar th er e as tw ar d, th is w as o ve rs te pp ed b y flo w s o f t he a ku ar ut u ni t ( 51 3) o f t he m al ig ât f or m at io n, w hi ch in va de d th e s ed im en ts (i n) b en ea th g ie se ck e m on um en t. bl ue th re edi gi t n um be rs ar e l ith ol og ic al co de s, ot he r a nn ot at io ns as in th e o rig in al . e xc er pt fr om th e s ou th n uu ss ua q se ct io n (p ed er se n et a l. 19 93 ). 87 ridge above tartunaq (south nuussuaq section at 70 to 76.5 km) there are semicontinuous exposures of this flow showing that it is invasive. it is particularly well exposed beneath point 975 m (figs 65, 66). the flow, which is compositionally quite similar to the invasive lava flows south of giesecke monument, has intruded within the top part of a mudstone overlain by a sandstone horizon. the mudstone belongs to the lower part of the assoq member (formerly the aussivik member). the overlying 100 m of sandstone belongs to the upper part of the assoq member (formerly the point 976 member of koch 1959). the upper contact of the flow is developed as a pillowed sill (fig. 76) similar to that beneath giesecke monument. koch (1959, p. 28–29) noted the curious combination of extrusive and intrusive features of the lavas around point 975 m, but the concept of invasive lava flows was not developed at the time. in the inland areas around the valley puiatussuaq north-east of the vaigat coast, this flow and an underlying, similar invasive flow of unit 512 can be followed as two prominent ledges in the sediment-dominated and scree-covered terrain. mf68_17cmimg_8071evsmodcut_c.eps 513 513 inv 513 inv 514 sed sed fig. 75. lava flows of the akuarut unit (513) and an intercalated flow of unit 514, invasive (inv) into mudstones (sed) of the assoq member. ridge just south of giesecke monument, south nuussuaq (south nuussuaq section at 58.9–59.3 km, see fig. 74). photo: erik vest sørensen. 8888 upper rinks dal member summary of the main features of the upper rinks dal member • comprises five units (514–518) defined by chemistry (intervals with variations in tio2 contents) but lithologically indistinguishable from each other; therefore shown as one group on maps and photogrammetric sections. • covered the flat surface of the akuarut unit on disko and part of southern nuussuaq and extended farther east onto the fluvial plain of the assoq member by onlap and invasion. stepped farther north onto the vaigat formation picrite shield. extends in remnants to the north coast of nuussuaq and onlaps gneisses on eastern nuussuaq. • eruption sites are known from west disko. • individual lava flows extend over large areas. flow volumes are up to 10 km3 but generally between 0.5 and 5 km3. • the uppermost unit (518) is retracted in distribution relative to the earlier units of the upper rinks dal member; it is also the most evolved unit and may reflect waning magma production in the final stage of the rinks dal member volcanism. the upper rinks dal member has been divided into five chemostratigraphic units (514–518). the division relies solely on chemical analyses (fig. 21), and distinction is not possible outside the analysed profiles. the geological maps and sections therefore show the upper rinks dal member as one undivided unit. the upper rinks dal member is present throughout disko and central to eastern nuussuaq from the south coast to the north coast. thicknesses are greatest, 500– 600 m, on western disko and decrease to around 400 m on central and southern disko and 300 m on northern and eastern disko. on nuussuaq, the upper rinks dal member is 340 m thick between paatuut and giesecke monument and decreases eastwards to 300 m. on central-eastern and north-eastern nuussuaq only the upper rinks dal member is present, and its thickness is reduced to c. 200 m due to onlap on the picrites of the vaigat formation and on the gneiss in the easternmost areas. around point 2000 m in the easternmost lava areas on nuussuaq, thicknesses locally reach 360 m due to ponding of flows in depressions on the old gneiss surface. the number of flows is largest on western disko, where there are around 25 flows present in any vertical succession, giving an average flow thickness of around 20 m, but with considerable variation as seen in the geological profiles and sections. the number of flows decreases and the average flow thickness increases eastwards, and at skarvefjeld on southern disko, there are 13 flows with an average thickness around 30 m. at qinngusaq on eastern disko, there are nine flows with an average thickness of 37 m. at giesecke monument on south-eastern nuussuaq, there are 13 flows with an average thickness of 26 m, and at nunavik on eastern nuussuaq there are four flows with an average thickness of 60 m. the five units (514–518) are present over most of the region, which shows that the eastwards decrease in the number of flows affected all units equally. unit 514 (upper transition flows) composition. these flows form a transition from the akuarut unit (513) into the main upper rinks dal member. the development with time towards lower tio2 contents was gradual (fig. 21), and the lower boundary of the upper rinks dal member is therefore not always well defined. the flows of unit 514 have 2.4–3.2 wt% tio2 and 5.4–7.6 wt% mgo, and in a plot of tio2 vs mgnumber they fall in the same field as the lower transition flows (unit 512) to which they are very similar. mf69_pillowedsill_ggurap69.eps black shale, assoq member 1 m zone of hydrothermal alteration brecciated shale and basalt basaltic sill (invasive lava flow) 1 m fig. 76. the upper contact between an invasive lava flow of the akuarut unit and mudstone of the assoq member. the flow forms a ‘pillowed sill’: irregular lava tongues intrude the mudstone in a zone that is more than a metre thick, and both mudstone and basalt are brecciated and hydrothermally altered at the contact. the relations show that the flow invaded the mudstone while this was still waterrich and only partly consolidated. south coast of nuussuaq, the slope beneath point 975 m (fig. 65). redrawn after pedersen (1975b). 89 petrography. unit 514 basalts are commonly almost aphyric with very scarce phenocrysts of plagioclase. the aphyric rocks range in tio2 between 2.4 and 3.02 wt%. some unit 514 basalts contain abundant microphenocrystic plagioclase and augite and scarce pseudomorphs after olivine, exemplified by fig. 77. basalts with a very finegrained groundmass and glomerocrystic plagioclase and pseudomorphed olivine and scarce augite also occur. distribution. the flows of unit 514 are found throughout disko and south-eastern nuussuaq (fig. 15). they form a 20–80 m thick succession of one to five flows, some of which may interdigitate with the upper part of the akuarut unit, as described above, and some interdigitate with unit 515, particularly on southern nuussuaq. unit 514 is absent in a few profiles. lithologies. descriptions are included in the lithology section on the entire upper rinks dal member. units 515, 516 and 517 (the main upper rinks dal member) composition. these three units, which constitute the main part of the upper rinks dal member, comprise basalts with tio2 contents within the rather wide range of 1.5–3.4 (3.7) wt%. the variation at any stratigraphic level is usually within one per cent tio2, typically in the interval 2–3 wt% tio2, but in all profiles there is an interval around the middle part with lower tio2 contents of 1.5–2.0 wt% (unit 516), which is used as a stratigraphic partitioning (fig. 21). the flows below and above this interval (units 515 and 517) are chemically indistinguishable and all three units are considered to be closely related. they are treated here under one heading. in a tio2 vs mg-number diagram (see fig. 97), most samples of units 515, 516 and 517 plot together with the lower rinks dal member and separate from the akuarut unit and the two units of transition flows. some flows of unit 515 show compositional overlap to unit 514, and the two units interdigitate on southern nuussuaq. petrography of unit 515. unit 515 is dominated by basalts with abundant plagioclase phenocrysts and plagioclaseaugite glomerocrysts up to 2 mm in size in a fine-grained groundmass. pseudomorphs after olivine less than 1 mm in size are scarce but present in all rocks (fig. 78a). the range of tio2 in this main type is 1.7–3.2 wt%. another common type has only scattered, 1 to 2 mm large plagioclase and augite glomerocrysts but carries abundant microphenocrystic plagioclase and augite less than 0.5 mm in size. pseudomorphs after microphenocrystic olivine are also present. a few basalts are almost aphyric and contain only scarce microphenocrysts and phenocrysts of plagioclase, augite and olivine pseudomorphs (fig. 78b). the range in tio2 for this type is 2.5–3.4 wt%. petrography of unit 516. the low-tio2 basalts of unit 516 (1.5–2.0 wt% tio2) are all strongly plagioclase and augite glomerophyric with 1–3 mm glomerocrysts. pseudomorphs after olivine are also present. no phenocrystpoor or aphyric samples were found. groundmasses range from very fine grained to almost doleritic. none of the samples are obvious plagioclase cumulates, and none are distinctive petrographical markers as compared to basalts of units 515 and 517. petrography of unit 517. the basalts of unit 517 show a considerable range in texture, varying from a few very phenocryst-poor basalts with fine-grained groundmass, through microphyric or micro-glomerophyric basalts mf70_8cm332813polcut_c.eps pl cpx 5 mmol fig. 77. thin section (scanned) of a lava of rinks dal member unit 514, upper transition flows. basalts with abundant microphenocrystic plagioclase (pl) and augite (cpx) and scarce olivine (ol) pseudomorphs. sample 332813, paatuut, southern nuussuaq. with very abundant microphenocrystic plagioclase and augite, to common, distinctly porphyritic basalts; these have plagioclase and augite glomerocrysts up to 5–10 mm in size, individual plagioclase phenocrysts up to 5 mm, augites up to 4 mm, and distinctive olivine phenocrysts (pseudomorphed) up to 2 mm (fig. 79). distribution. units 515 and 516 have the same distribution as unit 514 throughout disko and south-eastern nuussuaq and are not found east of the boundary fault on nuussuaq. unit 517 is the most widespread of the 12 units in the rinks dal member; it crossed the eastern boundary fault and reached northern nuussuaq where it onlapped the picrite shield of the vaigat formation and the high gneiss terrain in the east (fig. 15). it is present in sampled profiles at both saqqaq and nunavik (figs 11, 12). lithologies. descriptions are included in the lithology section on the entire upper rinks dal member. 5 mm mf71a_8cm354773polcut_c.eps a 5 mm mf71b_8cm332917upmodcut_c.eps b fig. 78. thin sections (scanned) of typical lavas from rinks dal member unit 515. a: basalt with abundant plagioclase phenocrysts and plagioclase-augite-olivine glomerocrysts in a fine-grained groundmass. plagioclase is colourless, augite (near lower edge) grey and olivine pseudomorphosed by brown and yellow clay. sample 354773, kvandalen, eastern disko. b: nearly aphyric basalt with scarce microphenocrysts and phenocrysts of plagioclase, augite and olivine pseudomorphs. sample 332917, kvandalen, eastern disko. mf72_8cm326430polmodcut_c.eps cpx pl 5 mm ol fig. 79. thin section (scanned) of a lava from rinks dal member unit 517. basalt with common plagioclase-augite glomerocrysts and individual phenocrysts of plagioclase (pl), augite (cpx) and distinctive olivine (ol, black pseudomorphs). sample 326430, marraat qaqqaat, south-eastern disko. 91 unit 518 (uppermost flows) composition. the uppermost flows of the rinks dal member have increased tio2 contents relative to the underlying succession (units 514–517), and more than half of them have just as high tio2 as the akuarut unit (fig. 21). in general, they have more than 2.5 wt% tio2 and mostly more than 2.8 wt%. in particular, one to two flows with up to 4.4 wt% tio2 occur in many profiles. this unit plots in a tio2 vs mg-number diagram in the same area as the akuarut unit and the transition flows; however, unit 518 and the akuarut unit can be distinguished by their trace elements (see below and larsen & pedersen 2009). petrography. the basalts of unit 518 show a considerable range in textures. slightly less than half of the rocks are phenocryst-poor or almost aphyric. many basalts are rich in plagioclase and augite microphenocrysts and glomerocrysts. one such lava flow contains decimetre-sized lava pegmatites (fig. 80a) which were dated by storey et al. (1998). finally, there are glomerophyric basalt samples with up to 6 mm large glomerocrysts of plagioclase, augite and pseudomorphs after olivine (fig. 80b). distribution. unit 518 is the only unit of the rinks dal member that has a more limited distribution than its predecessor (fig. 15). on western disko, it typically consists of four to five flows with a combined thickness of 100–150 m; on central and parts of eastern disko, there are one to two flows with a combined thickness between 20 m and 100 m. no flows reached north-eastern disko and nuussuaq. the unit is interpreted as representing a waning stage of the volcanism of the rinks dal member, when dwindling volumes of increasingly evolved magmas were erupted, centred on western disko or west of disko. lithologies. descriptions are included in the lithology section on the entire upper rinks dal member. lithologies of the upper rinks dal member all lava flows in the upper rinks dal member are subaerial. on easternmost disko and south-easternmost nuussuaq, they are also invasive into, or flowed onto, wet sediments, with ensuing development of colonnades and entablatures in the flows. however, at this stage the assoq lake was reduced to small temporary pools, coal swamps and fluvial plains. on north-easternmost nuussuaq, the flows of unit 517 ponded in water-filled depressions on the gneiss surface, leading to development of thick flows with prominent entablatures. 10 mm mf73a_8cm328406upcut_c.eps a 5 mm mf73b_8cm327016polmodcut_c.eps b fig. 80. thin sections (scanned) of rinks dal member unit 518. a: pegmatitic vein with large crystals of plagioclase (white) and augite (grey) and smaller crystals of ilmenite and magnetite (black). the matrix of the same phases shows micrographic intergrowths. dated sample 328406 (storey et al. 1998), daugaard-jensen dal, southcentral disko. b: sparsely glomerophyric basalt with up to 6 mm large glomerocrysts of plagioclase and olivine (pseudomorphosed) in a very fine-grained groundmass. sample 327016, skarvefjeld, south disko. 9292 on western disko, the upper rinks dal member displays a significant variation in lava flow thickness and morphology, whereas these are more uniform on central and eastern disko and nuussuaq. the few eruption sites that have been found are all located on western disko. together with the higher number of flows and larger total thicknesses in the west, these features support the interpretation that the primary source areas of these basalts were on western disko and west of disko. eruption sites on western disko uiffaq, unit 515. a volcanic eruption site is exposed in the steep coastal cliff on the uiffaq peninsula c. 14 km west of the disko gneiss ridge at nuuk kangilleq (fig. 81). here a brecciated basaltic feeder dyke can be followed vertically for about 180 m until it ends in a funnelshaped, reddish scoria zone within a basaltic lava flow (fig. 82) situated in the middle of nine successive lava flows of unit 515. the lower part of the feeder dyke is strongly brecciated and crudely bedded (fig. 83) and was described briefly by pedersen (1977c, p. 63) who suggested that the dyke was brecciated when it came into contact with a local aquifer. nordfjord, unit 517. a prominent basaltic volcanic neck is exposed high in the steep mountain wall along the south side of nordfjord/kangersooq c. 2 km wnw of point 1137 m (see fig. 103, loc. 25). it was originally mapped as part of the nordfjord member by pedersen & ulff-møller (1987). a later visit revealed that the neck is an eruption site feeding basalts of unit 517 of the upper rinks dal member. the semi-cylindrical neck has a diameter of c. 500 m and comprises two basalt bodies, of which the upper one is part of a lava flow originating from the neck (fig. 84). the top of the neck is poorly exposed except in the near-vertical northern wall, but a more than 50 m thick coarsely columnar-jointed lava flow of unit 517 can be seen south-west of the neck at a distance of c. 500 m (fig. 85). two sample profiles, at distances of c. 500 m and 900 m from the neck, through the top of the upper rinks dal member and the nordfjord member (fig. 86) demonstrate that at the time of eruption of the nordfjord member (unit 520), the neck and its toplava constituted a prominent hill which allowed airborne tuffs to settle on it, while a total of 85 m of andesitic and dacitic lava flows were banked up against its flank. the neck constitutes the largest known basaltic eruption site of the maligât formation. lithologies on disko west of the disko gneiss ridge. the general lithologies are well illustrated by the section along the north coast of kangerluk between naqerloq and illukasik (south disko section at 18–42 km). the upper rinks dal member is here c. 520 m thick and consists of about 25 lava flows (fig. 8, profile 1, and fig. 87). west of tunup qaqqaa, 51 52 53 kmfig. 82-83 400 600 800 m m7 m8.25 m8 m8.32 m8.26 no1 uiffaq iron locallity nuuk kangilleq δc δc fe 520 514–518 514–518 513 fe cr gv03_03_055_lml nw se fig. 81. photogrammetrically measured section of a part of the south-west-coast of the uiffaq peninsula, south disko. a feeder dyke and small eruption crater (cr) in the upper rinks dal member are located at 52.9 km (shown in figs 82, 83). a native-iron-bearing dyke (red) described later cuts obliquely through the exposure at 50–51 km and appears at the coast at 51.2 km in the famous locality where native iron was first described (nordenskiöld 1871). blue three-digit numbers are lithological codes; other annotations as in the original. excerpt from the south disko section (pedersen et al. 2003). 93 the thickness variation of the flows is particularly large, as illustrated by the south disko section at 23–30 km, where individual lava flows range in thickness from less than 10 m to 60 m (the thick marker flow m8.11) and there are intercalated flow fields of thin pahoehoe flows (e.g. unit m8.13) that can be followed over more than 14 km and which were at least partially erupted from a crater site near ukaleqartarfik (fig. 88). fig. 82. eruption site for a lava flow of unit 515, nuuk kangilleq, uiffaq (fig. 81). the brecciated feeder dyke can be followed vertically for about 180 m until it ends in a funnel-shaped reddish scoria zone within a lava flow. photo: finn ulff-møller. mf75_8cmfum_1984_02_14modcut_c.eps 514 515 515 513 mf76_8cmakp_1974_uiffacbrdyke1udglmodcut_c.epsfig. 83. the lower part of the brecciated feeder dyke shown in fig. 82. height of field of view 8–10 m. fig. 84. eruption site for rinks dal member unit 517 on the south side of nordfjord, west disko (for locality, see fig. 103, loc. 25). the dark exposure faces north; yellow lines schematically trace neck outline, base of thick lava flow and columnar jointing. the semicylindrical neck has a diameter of c. 500 m. it is composed of a lower body of compact basalt showing coarse columnar jointing roughly perpendicular to the wall, and an upper basalt body with a compact lower part with vertical columns up to several metres thick and an upper vesiculated part. the upper body is part of a lava flow originating from the neck (fig. 85). mf77_11cmakp_34_c3_79modcut_c.eps 9494 mf78_17cmakp_1985_08_33l85-28modcut_c.eps 520 520 sed+tuff 517 530 517 lava flow from neck fig. 85. the lava succession c. 500 m south-west of the volcanic neck in fig. 84. the more than 50 m thick flow of unit 517 was erupted from the neck; it is the third flow beneath the top of the rinks dal member. the flows of unit 517 are covered by sediments, including tuffs, and basalt lava flows of the nordfjord member (unit 520), followed by flows of the niaqussat member (unit 530). 1050 1000 950 900 m 1050 1000 950 900 900 m 800 700 600 500 400 ? 300 older lava flows 200 100 centre of neck volcanic neck ew 0 m 279060 279059 279058 basalt tuff 279057 279056 279300 279299 279298 279297 279278 279295 279296 279294 laterite tuff and claystone 279292-93 baked acid tuff 279073 lava flow fed from neck 279072 tuff, soil rhyolite tuff 279074 1000 900 m 1000 900 900 m 800 700 600 500 400 300 200 100 no vertical exaggeration niaqussat member nordfjord member niaqussat member nordfjord member nordfjord member rinks dal member tuffs tuffs tuffs rinks dal member rinks dal member rinks dal member ew 0 ? m gv04_06_029_lml 518 518 517 517 a b fig. 86. eruption site for rinks dal member unit 517 on the south side of nordfjord (figs 84, 103). a: schematic reconstruction. b: profiles through the neck and the surrounding lava succession. for colours and symbols, see fig. 7. see text for interpretation. 95 fig. 87. the full upper rinks dal member, 520 m thick, on south-central disko. the boundaries between the successive units in the upper rinks dal member are not precisely located because the tunup qaqqaa sample profile was not flow-by-flow sampled (fig 8, profile 1). a: photograph. b: photogrammetric interpretation, excerpt from the south disko section (pedersen et al. 2003). some characteristic marker flows are separately coloured and annotated. the rinks dal member is overlain by the nordfjord member (unit 520, red and pink) and the niaqussat member (unit 530, olive green). note the red scoria of the eruption crater (cr) for the lowest nordfjord member flow. south side of the tunup qaqqaa mountain, south-central disko. 40 km39 1000 m 800 600 400 200 m4 m7 m8.3 m7 m8.18 m8.20 m8 m8.22 m8.19 ni3 no1 δm crcr tunup qaqqaa gv03_03_056_lml w e b mf80a_17cml9118_3_02mod_c.eps 520cr 513 514–518 530 a 9696 mf81a_17cma9122_3_02mod_c.eps 520 a 514–518 25 km24 200 400 600 800 m m8 m8.12 m8 m8.16 m8.4 m8.13 m8.5 m8.10 no2 no1 m cr ukaleqartarfik gv03_03_057_lml nnw sse b fig. 88. the upper rinks dal member on south-west disko. a: photograph. b: photogrammetric interpretation, excerpt from the south disko section (pedersen et al. 2003). some characteristic marker flows are separately coloured and annotated. note the intercalated thick and thin flows and the eruption crater (cr) for the flow field of thin flows of marker unit m8.13, marked with a black star in a. the boundaries between the successive units in the upper rinks dal member cannot be located because no sample profiles are close enough. flows of the nordfjord member (unit 520) overlie the rinks dal member. ukaleqartarfik 15 km west of tunup qaqqaa, south-west disko. 97 fig. 89. skarvefjeld on the south coast of disko, exposing a large part of the rinks dal member and the upper boundary to the nordfjord member. some characteristic marker flows are separately coloured and annotated. blue three-digit numbers are lithological codes; other annotations as in the original. the photo in fig. 56 covers the section from 85.3 to 86.7 km on the base line. excerpt from the south disko section (pedersen et al. 2003). 85 86 km84 800 1000 m 200 400 m8 m6 m8 no1 m3.1 m3.1 m6 m4 s lsls m7 m8.17 m8.17 m7.1 m5 m3 m8.17 innarsuaq innap qaqqaa qeqertaluk 509 511 511 513 514 515 517 518 520 skarvefjeld gv03_03_058_lml wsw ene mf83_17cmimg_9424evsmod_c.eps 515 515 (m8.17) 514 ? 513513 514 ? 517 518 fig. 90. lava flows of the upper rinks dal member at innap qaqqaa, the highest part of the skarvefjeld mountain. compare with the photogrammetric interpretation in fig. 89 where the flow labelled m8.17 is identified. photo: erik vest sørensen. 9898 central and southern disko. on the disko gneiss ridge and just east of it, the upper rinks dal member forms a fairly monotonous succession of subaerial lava flows with blocky pahoehoe flow morphology and with only very limited lithological variation. the lava succession is well illustrated by the coastal wall at skarvefjeld between blæsedalen and brededal on the south coast of disko, shown in fig. 89 and fig. 7, profile 4. the upper rinks dal member is here 390 m thick and consists of 13 subaerial lava flows, which vary in thickness between 10 m and c. 50 m. (fig. 90). some of the flows have slightly eroded top surfaces with a few centimetres of lateritic soil while others are devoid of soil. the lava flows show a crude columnar jointing with no indication of the presence of water. eastern disko. on eastern disko, the upper rinks dal member thins gradually eastwards from 300 m to less than 200 m, and the number of lava flows decreases to less than ten. individual lava flows more than 40 m thick occur, e.g. fig. 7, profile 10 (skorstensfjeld) or profile 12 (point 1123 in kvandalen). the lower part of the succession shows eastwards increasing interaction with unconsolidated sand and coal swamps of the upper part of the assoq member. the flows ponded in topographic lows and are strongly columnar-jointed as illustrated by the two lowermost flows of the upper rinks dal member exposed in a corrie in blåbærdalen 2 km north-west of skorstensfjeld (fig. 91). the upper of these two flows is capped by a prominent horizon of red soil. an example of a very thick, ponded flow, which reaches 90 m in thickness is seen at aqajaruata qaqqaa (fig. 7, profile 13). well developed, curved colonnades and prominent entablature zones attest to ponding in a water-filled depression. the uppermost part of the rinks dal member is particularly well illustrated by the basalt plateau exposed at point 1014 m in sortebærdalen (fig. 7, profile 11). in the lower part of the succession, three lava flows of units 512−514 have invaded friable sandstone with a thick coal bed referred to the assoq member. the overlying six lava flows of units 515−518 have no traces of sandstone between them (fig. 92), suggesting that at this stage the assoq lake basin was obliterated throughout the disko area. mf84_17cmimg_9303evsmodcut_c.eps 513 inv 514 515 100 m assoq mb assoq mbponded flow laterite lat. fig. 91. the middle and upper rinks dal member on eastern disko. a single flow of the akuarut unit (513 inv) is invasive into sediments of the assoq member. the following two flows of the upper rinks dal member (unit 514) are subaerial and ponded in wet depressions where they developed pronounced columnar jointing; the upper flow is capped by a prominent laterite horizon (lat). the overlying two flows (unit 515) are purely subaerial. north-exposed corrie wall in blåbærdalen 2 km north-west of skorstensfjeld, eastern disko. photo: erik vest sørensen. 99 lithologies on nuussuaq south coast of nuussuaq. the original thickness of the upper rinks dal member can only be assessed at a few localities where the overlying nordfjord member is preserved. along the south coast, the upper rinks dal member is c. 300 m thick at point 1888 m north of ataata kuua (fig. 11, profile 2) where it consists of c. 15 flows. farther to the east the upper rinks dal member is c. 340 m thick at point 2010 m (fig. 93) and at giesecke monument where it consists of 12–13 flows (figs 74, 94). the flow thickness varies from less than 10 m to a maximum of c. 70 m for the large flow mm in the south nuussuaq section (fig. 93). this flow may have a local origin (see map fig. 96), but most of the other flows are considered to have flowed in from the south-west, partly channelled along the south-easterly sloping surface of the vaigat formation picrite lava shield. the easternmost exposure of the upper rinks dal member is a just 40 m thick succession of three subaerial lava flows of unit 517, which have crossed the eastern boundary fault to be emplaced on gneiss north of saqqaq (pedersen & larsen 1987; pedersen et al. 2007a; fig. 11, profile 11) where they are covered by lava flows of the nordfjord member and separated from these by traces of soil. nunavik ridge. north of the aaffarsuaq valley, lava flows of the upper rinks dal member unit 517 rest either on picrite lavas from the vaigat formation or on gneisses. they cap a number of local high mountaintops and are overlain by lavas of the nordfjord and niaqussat members. they are best illustrated by their occurrence along the mountain ridge nunavik (nuugajukassak alleq and qulleq; larsen & pedersen 1992, figs 3 and 4, central fig. 92. the rinks dal member on eastern disko. single flows of units 512 and 513 are invasive into assoq member sandstones. the overlying upper rinks dal member comprises a unit 514 flow invasive into coal-bearing assoq member sands, followed by six subaerial flows of units 515–518 without traces of sandstone between them. the upper part of sample profile 11 in fig. 7 is situated just left of the ridge in the centre of the image, and the lower part is situated to the right of the ridge. the level of a poorly exposed coal layer more than 1.5 m thick is indicated (coal). height of field of view 300 m. basalt plateau at point 1014 m in sortebærdalen, eastern disko, seen from the south-west. photo: erik vest sørensen. mf85_17cmimg_0529evsmod_c.eps 514 inv 512 inv 513 inv assoq mb 518 516 515 coal 100100 gv03_03_059_lml 45 46 47 km 48 200 400 600 800 1000 1400 1600 1800 2000 m j 450 450 450 520 513 512 458 450 514–517 im i2 i3 ml ml i3 i2 ml mm mn mt i1 point 2010 m naujât mb level not exposed atane fm qilakitsoq mb eruption site sills nw se mf87_17cmimg_8058evsmod_c.eps 520 ? 517 516 515 514 513 101 nuussuaq section at 59 to 71 km, units mu and mm; fig. 163, unit mf2). they are shown in fig. 12, profile 13 and fig. 95, where they form a succession of about seven lava flows with a total thickness of up to 315 m. some of the flows are ponded and as much as 75 m thick, and one flow locally reaches a thickness of 90 m. several flows have prominent colonnades and entablatures suggesting solidification in a wet environment. lava volumes in the upper rinks dal member despite the fairly monotonous character of the upper rinks dal member, photogrammetric studies have identified a number of individual lava flows or flow fields. these are indicated with colours and/or annotation on the photogrammetric sections where they can be followed over distances ranging from a few kilometres to more than 15 km, commonly over 4–8 km (table 3). these measurements are not easy to convert into areas, but if the lengths are simply squared, and if a typical thickness of c. 30 m is assumed, many flows will have volumes between 0.5 and 3 km3, and very few will exceed 5 km3. the largest of the measured flows belongs to unit 514 and can be followed for a distance of 25 km along the south coast of nuussuaq from west of ataata kuua to giesecke monument. it is shown as marker unit mm in the south nuussuaq section (fig. 93). the thickness of this flow was measured by pedersen & dueholm (1992, fig. 13), who found that it varies from less than 45 m to more than 70 m within a ‘measurable’ area of 18 × 3–4 km (fig. 96). the flow has a minimum area of c. 240 km2 and its minimum volume can be estimated at 10 km3. the thickness measurements can be contoured and show a maximum near paatuut, suggesting either ponding or that the feeder zone might have been located in this part of nuussuaq; no actual feeders have been observed. the real volume of the flow may substantially exceed 10 km3. by comparison, the most voluminous lava flow known in the maligât formation is a native-iron-bearing composite lava flow of the nordfjord member in the mellemfjord area on south-western disko which has a volume of at least 14 km3 (pedersen 1977b, fig. 19 and page 41). facing page (top): fig. 93. photogrammetrically measured section along the south coast of nuussuaq, showing subaerial lava flows of the rinks dal member (units 512–517) and the nordfjord member (unit 520) overlying the vaigat formation (ordlingassoq member, units 450 and 458). blue three-digit numbers are lithological codes; other annotations as in the original. the synvolcanic mudstones of the naujât member are not exposed; their approximate level is indicated. point 2010 m near paatuut, excerpt from the south nuussuaq section (pedersen et al. 1993). facing page (bottom): fig. 94. giesecke monument, a landmark peak reaching 1574 m a.s.l. on the south coast of nuussuaq. subaerial flows of the upper rinks dal member (units 514–517) constitute the major part of the peak, which is topped by a remnant of a nordfjord member basalt flow (unit 520). the exposed section is 550 m thick. see also fig. 74. photo: erik vest sørensen. 102102 mf88_17cmakp_1991_10_07mod_c.eps 517 vf 520 530+532 vf gneiss fig. 95. the nunavik mountain ridge on northern nuussuaq, just east of the eastern boundary fault. here the gneiss rises to 1300–1500 m altitude and is overstepped by the vaigat formation (vf) in lows, overlain by seven flows of unit 517 of the upper rinks dal member, two basalt flows of the nordfjord member (520) and a series of picrite and basalt flows of the lower and upper niaqussat member (530, 532). unit 517 is here 315 m thick. the ridge is located in figs 4, 6. fig. 96. thickness map of the most voluminous lava flow measured in the upper rinks dal member. the thickness measurement points, tracing the coastal wall, are spaced over a sufficiently wide coastal strip that a contoured thickness map can be made. the flow has a minimum area of c. 240 km2 and a minimum volume of 10 km3. paatuut, south coast of nuussuaq. after pedersen & dueholm (1992, fig. 13). 70°18’n 52°35’w 0 60 60 50 45 m 50 70 70 n 5 km vaigat paatuut 70–80 m (>70 m) thickness of lava low at aa ta ku ua 60–70 m 50–60 m 40–50 m thickness measurement by multi-model photogrammetry gv01_02_163_lml 103 annotation localities length, km western end eastern end akuarut unit (513) south disko section (west to east) m 7.2 naqerloq to tunup qaqqaa 7.4 scree scree m 7.3 kangaarsuk to akuarut 12.5 air air m 7.1 skarvefjeld to ippik 3.3 scree air central disko section (west to east) m 7.1 perlertut qaqqaat 2.4 scree scree m 7.2 perlertut qaqqaat 1.6 scree scree m 7.3 perlertut qaqqaat 2.0 scree scree m 7.4 perlertut kuuat to stordal 3.0 air scree m 7.5 perlertut kuuat to stordal 2.9 air scree m 7.6 perlertut kuuat to stordal 2.8 air air m 7.7 stordal 8.5 featheredge featheredge m 7.8 stordal 6 featheredge featheredge m 7.9 stordal 3.9 featheredge scree m 7.10 kuugannguaq 2.8 scree scree m 7.11 narsap qaqqaa 5.0 scree scree south nuussuaq section m umiusat to tartunaq c. 12 air air median length3.3 km; mean length 5.1 km. upper rinks dal member (units 514–518) south disko section (west to east) m 8.8 point 639 m to nuuk qiterleq 6.1 scree air m 8.7 point 639 m to point 691 m 7.0 scree air m 8.15 point 639 m to point 691 m 6.6 scree air m 8.6 nuuk killeq to point 691 m 4.5 fault air m 8.14 point 691 m 2.6 air air m 8.5 nuuk killeq to umiiarneq 13.1 scree air m 8.13 nuuk killeq to umiiarneq 13.8 scree air m 8.4 nuuk kangilleq to naqerloq 8.9 scree air m 8.2 point 691 m 2.6 scree air m 8.12 ukaleqartarfik to umiiarneq 5.6 featheredge featheredge m 8.11 ukaleqartarfik to kuussuaq 8.0 featheredge near featheredge m 8.10 ukaleqartarfik to naqerloq 3.2 scree air m 8.1 naqerloq to kuussuaq 5.1 air air m 8.9 naqerloq to kuussuaq 7.7 scree featheredge m 8.21 kuussuaq 1.6 featheredge air m 8.24 tunup qaqqaa 3.5 air air m 8.22 tunup qaqqaa 3.5 air air m 8.20 tunup qaqqaa 4.9 featheredge air m 8.19 tunup qaqqaa 6.0 featheredge air m 8.18 tunup qaqqaa 7.0 featheredge air m 8.3 tunup qaqqaa 6.0 featheredge scree m 8.32 kangerluk to itilleq (uiffaq) 13.8 air air m 8.23 kangerluk to uiffaq 5.0 air featheredge m 8.27 uiffaq 5.0 air featheredge m 8.26 uiffaq 5.8 air featheredge m 8.25 uiffaq 10.0 featheredge scree m 8.28 uiffaq to itilleq 6.0 featheredge air m 8.29 killiit to lyngmarksfjeld 11.8 air air m 8.17 skarvefjeld to ippik 6.8 air air m 8.30 brededal to nuugaarsunnguaq 4.7 air featheredge m 8.31 brededal to niuluut 2.1 air scree central disko section (west to east) m 8.2 perlertut qaqqaat 4.3 fault air m 8.3 perlertut qaqqaat 2.8 scree scree m 8.6 perlertut qaqqaat to stordal 5.2 featheredge air m 8.8 stordal 6.3 air featheredge m 8.9 stordal 3.1 featheredge air m 8.12 kuugannguaq 4.3 air featheredge m 8.16 point 1530 m 5.7 air air m 8.14 point 1530 m to qinngusaq 17.5 featheredge air m 8.17 point 1530 m 6.0 panel boundary air m 8.15 narsap qaqqaa to qinngusaq 10.0 featheredge featheredge m 8.18 qinngusaq 3.0 panel boundary scree m 8.19 kvandalen 4.8 featheredge scree m 8.20 kvandalen 12.5 air scree south nuussuaq section m m point 1722 m to upalluk 25.6 air featheredge central nuussuaq section m m point 2080 m to point 2000 m 12.9 air air median length 5.9 km; mean length 7.0 km. note: similarly annotated flows in different sections are not identical flows. the annotation numbers were made during compilation and are not systematic with height or longitude. ends noted as scree or air are ‘open’, i.e. the flow may be longer. featheredges are true ends. table 3. measured minimum lengths of individual lava flows in the rinks dal member 104104 chemical compositions of the rinks dal member as stated earlier, the rinks dal member almost exclusively consists of three-phase cotectic basalts evolved in deepseated magma chambers (larsen & pedersen 2009). the crystallisation sequence is olivine (+ rare chromite) → olivine + plagioclase → olivine + plagioclase + clinopyroxene. fe-ti oxides are late. most lava samples contain small amounts of interstitial clay, probably altered glass. however, the samples are generally fresh, with an average volatile content of 1.27 ± 0.70 wt%; of 1011 chemically analysed samples only 14 have more than 4 wt% volatiles. mobile elements such as k, rb, and ba do, however, show some scatter. representative analyses are shown in table 4. in contrast to the vaigat formation, mgo is a poor variation parameter for the rinks dal member because of its narrow variation interval (mainly 4.4–9.2 wt% mgo; only 13 samples have higher mgo up to 25.4 wt%). the mg-number (mg = atomic 100mg/(mg + fe2+)) provides a much better data spread (mainly mg 35–64, but up to 83) and has therefore been chosen as the main variation parameter for the rinks dal member. this means that the geochemical plots for the vaigat and maligât formations shown in pedersen et al. (2017), and in the present work, are not directly comparable. for plots of the maligât formation with mgo as the common variation parameter, see larsen & pedersen (2009). the 12 chemostratigraphic units of the rinks dal member are difficult to show individually in chemical plots. therefore, in many plots the analyses have been grouped into the lower rinks dal member (units 505– 509, 512), the skarvefjeld unit (511), the akuarut unit (513), and the upper rinks dal member (units 514– 518). major elements variation diagrams for the major elements are shown in fig. 97. crystallisation proceeds from right to left in the diagrams. only the most mg-rich samples have crystallised olivine alone. plagioclase began crystallising around mg-number 58 (mgo around 7.5 wt%) where al2o3 began to decrease. it was followed shortly by clinopyroxene, and cao began to decrease also. na2o increased slightly but was also removed with the plagioclase. feo* increased markedly, and so did tio2, k2o and p2o5. the most fractionated basalts, with mg-number <40 (mgo <c. 4.8 wt%), may perhaps have started to fractionate titanomagnetite, whereas k2o and p2o5 remained incompatible throughout. the k2o diagram shows a fair amount of scattering due to secondary alteration. the sio2 trend forms a broad arc with maximum values that generally do not exceed 50.3 wt% sio2. of the analysed samples, 361 (36%) are q normative, but only 120 (12%) have >1% q. some flows with around and above 50 wt% sio2 may be slightly crustally contaminated. two flows with 51.3–51.4 wt% sio2 are clearly contaminated. the skarvefjeld unit does not follow the general fractionation trend shown by the other units. its major element trends have higher tio2, feo* and p2o5 and lower cao and al2o3 than the other units at similar mg-numbers; however, the trends merge into the main trend at low mg-numbers. as pointed out by larsen & pedersen (2009), the skarvefjeld unit is best explained as a result of an episode of mixing followed directly by eruption: a batch of picrite magma, perhaps unusually large, entered the deep-seated magma chamber and mixed with the residing fractionated melt there; the mixed magma was then erupted without further equilibration and fractionation. the skarvefjeld trends in fig. 97 are thus not fractionation trends but mixing curves. curves calculated by mixing of average unit 512 basalt with 6.2 wt% mgo (the residing melt in the chamber at the time) with a calculated unit 511 picrite with c. 21 wt% mgo model the skarvefjeld trends very well (fig. 97). there is also petrographic support for this explanation because some of the samples with 12−13 wt% mgo contain plagioclase glomerocrysts although plagioclase is not on the liquidus in such mgo-rich magmas. three samples with 18−24 wt% mgo (mg-number >75) contain accumulated olivine; plagioclase phenocrysts were not observed in these. trace elements variation diagrams for trace elements are shown in figs 98, 99. the incompatible elements (fig. 98) show a general increase with decreasing mg-number, but significant scatter is apparent for some elements. this is primarily due to the fact that the skarvefjeld unit and the upper and lower transition units comprise some lavas that are significantly enriched in rb, ba, sr, nb, ce, and zr/y relative to the ‘normal’ lavas. these enriched lavas are only present in a small area on north-eastern disko and on southern nuussuaq (figs 10–11). it is also notable that most flows of the akuarut unit have particularly high contents of sr and high zr/y. of the transition elements (fig. 99), v, cu, and zn increase throughout the sequence of melt evolution. 105 40 45 50 55 60 30 35 40 45 50 55 60 65 70 75 80 85 upper rinks dal mb (units 514-518) akuarut unit (513) skarvefjeld unit (511) lower rinks dal mb (units 505-509, 512) 5 10 15 20 30 35 40 45 50 55 60 65 70 75 80 85 5 10 15 20 30 35 40 45 50 55 60 65 70 75 80 85 0 5 10 15 30 35 40 45 50 55 60 65 70 75 80 85 0 1 2 3 4 5 30 35 40 45 50 55 60 65 70 75 80 85 0 1 2 3 4 5 30 35 40 45 50 55 60 65 70 75 80 85 0.0 0.5 1.0 1.5 30 35 40 45 50 55 60 65 70 75 80 85 0.0 0.2 0.4 0.6 0.8 30 35 40 45 50 55 60 65 70 75 80 85 sio2 al2o3 feo* cao tio2 na2o k2o p2o5 mg-number mg-number fig. 97. major element variation diagrams for rocks of the rinks dal member. data in wt% oxides recalculated volatile-free. feo* is total iron as feo. the curves in all diagrams are mixing curves calculated by mixing of an average unit 512 basalt with a modelled unit 511 picrite composition (tick marks shown for every 10% increment). the mixing trend reproduces the skarvefjeld trend well; see text for discussion. 106106 0 10 20 30 40 30 35 40 45 50 55 60 65 70 75 80 85 0 50 100 150 200 250 300 30 35 40 45 50 55 60 65 70 75 80 85 0 100 200 300 400 30 35 40 45 50 55 60 65 70 75 80 85 0 10 20 30 40 50 60 70 30 35 40 45 50 55 60 65 70 75 80 85 0 100 200 300 400 30 35 40 45 50 55 60 65 70 75 80 85 0 5 10 15 20 25 30 35 30 35 40 45 50 55 60 65 70 75 80 85 0 1 2 3 4 5 6 7 30 35 40 45 50 55 60 65 70 75 80 85 mg-number mg-number 0 10 20 30 40 50 60 70 80 90 30 35 40 45 50 55 60 65 70 75 80 85 rb ba sr y zr nb zr/y ce uppermost flows (518) upper rinks dal mb (515–517) enriched upper transition flows upper transition flows (514) akuarut unit (513) lower transition flows (512) enriched skarvefjeld skarvefjeld unit (511) lower rinks dal mb (505–509) fig. 98. incompatible trace-element variation diagrams for rocks of the rinks dal member. data in ppm. 107 mg-number mg-number 0 100 200 300 400 500 600 30 35 40 45 50 55 60 65 70 75 80 85 0 10 20 30 40 50 30 35 40 45 50 55 60 65 70 75 80 85 0 100 200 300 400 500 30 35 40 45 50 55 60 65 70 75 80 85 0 50 100 150 200 30 35 40 45 50 55 60 65 70 75 80 85 0 100 200 300 400 30 35 40 45 50 55 60 65 70 75 80 85 0 200 400 600 800 30 35 40 45 50 55 60 65 70 75 80 85 three samples off scale at 900-1450 ppm three samples off scale at 900-1450 ppm uppermost flows (518) upper rinks dal mb (515–517) enriched upper transition flows upper transition flows (514) akuarut unit (513) lower transition flows (512) enriched skarvefjeld skarvefjeld unit (511) lower rinks dal mb (505–509) v sc cu zn ni cr fig. 99. transition-element variation diagrams for rocks of the rinks dal member. data in ppm. 108108 sc peaks around mg-number 50–55 and is then lowered by fractionation of clinopyroxene. ni and cr are compatible in olivine, chromite and clinopyroxene and decrease throughout the sequence of melt evolution. the skarvefjeld unit has relatively low sc and high ni and cr which is an effect of magma mixing, as described above. ree and multi-element diagrams are shown in fig. 100. with few exceptions, the ree patterns are parallel and show hump-backed shapes with (la/sm)n <1 and (tb/lu)n >1. parallel displacement of patterns correlate well with mg-numbers and is due to ordinary olivine plagioclase-augite fractionation. the low-mg, evolved rocks of the akuarut unit show small eu troughs, presumably due to extended plagioclase fractionation. the ree patterns of the akuarut unit are steeper than those of the other units and are tilted relative to these, with higher lree and lower or equal hree. this can be modelled by involving stabilisation and fractionation of small amounts of garnet (c. 1%) in the deep-seated magma chamber at this stage (larsen & pedersen 2009), which will also explain the high zr/y. the multi-element patterns have curved shapes like the ree patterns, with relative depletion of the most incompatible elements rb–ba–th–u, and distinct troughs for k, pb and p. there are some variations in the levels of rb and k of individual samples which may be due to secondary alteration. the increased levels of the most incompatible elements in the enriched samples fig. 100 (first of two parts). ree and multi-element diagrams for representative rocks of the rinks dal member. left diagram: chondrite normalised; right diagram: primitive mantle normalised; normalisation factors from mcdonough & sun (1995). disko-1 from unit 509 is geus’ in-house geochemical standard. 1 10 100 1000 la ce pr nd sm eu gd tb dy ho er tm yb lu lower rinks dal member units 505, 506, 509 unit 507 (low ti) disko-1 1 10 100 1000 rb ba th u nb ta k la ce pb pr sr nd p sm zr hf eu ti gd tb dy y yb 1 10 100 1000 la ce pr nd sm eu gd tb dy ho er tm yb lu skarvefjeld unit unit 511 enriched 1 10 100 1000 rb ba th u nb ta k la ce pb pr sr nd p sm zr hf eu ti gd tb dy y yb 1 10 100 1000 a la ce pr nd sm eu gd tb dy ho er tm yb lu akuarut unit and lower transition flows akuarut unit (513) lower transition flows (512) 1 10 100 1000 rb ba th u nb ta k la ce pb pr sr nd p sm zr hf eu ti gd tb dy y yb 109 shown in fig. 98 are clearly seen in fig. 100 as well. in the most enriched samples, such as 136969 of the skarvefjeld unit and 318758 of the upper transition flows (table 4), the increase comprises all elements from rb to ti but not the hree. these enriched flows bear an imprint of accidental contamination with an enriched mantle component, as discussed in larsen & pedersen (2009). sample 327065 of the lower transition flows (table 4) has a different enrichment pattern with a small increase in the most incompatible elements from rb to pb but not other elements; the sr–nd isotope analyses suggest that this increase is caused by slight crustal contamination. 1 10 100 1000 la ce pr nd sm eu gd tb dy ho er tm yb lu upper transition flows unit 514 enriched 1 10 100 1000 rb ba th u nb ta k la ce pb pr sr nd p sm zr hf eu ti gd tb dy y yb 1 10 100 1000 la ce pr nd sm eu gd tb dy ho er tm yb lu rdm uppermost flows unit 518 1 10 100 1000 rb ba th u nb ta k la ce pb pr sr nd p sm zr hf eu ti gd tb dy y yb 1 10 100 1000 la ce pr nd sm eu gd tb dy ho er tm yb lu main upper rinks dal member units 515, 517 unit 516 (low ti) 1 10 100 1000 rb ba th u nb ta k la ce pb pr sr nd p sm zr hf eu ti gd tb dy y yb b fig. 100 (second of two parts). ree and multi-element diagrams for representative rocks of the rinks dal member. left diagram: chondrite normalised; right diagram: primitive mantle normalised; normalisation factors from mcdonough & sun (1995). 5052 5061 5071 5091 5092 5111 5111 5112 5111 5121 5121 5131 354719 332881 326406 340961 157246 136969 332930 362348 136973 362344 327065 279041 5324.25 5319.33 5338.17 5342.09 5314.41 5324.06 5318.62 5215.44 5324.11 5215.40 5319.79 5307.77 6935.355 6958.587 6916.251 6915.768 6918.378 7009.380 6958.831 6939.040 7009.392 6939.014 6917.262 7005.896 18 786 162 103 148 1256 1184 639 1207 687 312 1193 48.43 48.72 48.24 49.57 49.10 44.82 47.88 48.21 48.17 48.93 48.88 49.13 2.94 1.88 1.66 2.00 2.15 1.36 2.10 2.16 2.41 2.94 2.69 3.44 13.69 14.59 15.26 14.48 13.91 11.01 13.27 13.92 13.85 13.64 13.12 13.12 5.25 4.70 4.58 1.73 3.70 2.08 5.54 2.25 4.92 3.05 4.80 4.53 8.62 6.76 6.91 10.20 8.90 8.50 6.24 9.49 7.39 9.94 8.67 9.25 0.24 0.18 0.20 0.18 0.20 0.17 0.18 0.18 0.19 0.19 0.21 0.21 5.99 8.06 6.92 6.86 6.47 18.09 10.39 8.07 6.65 6.33 5.83 5.76 10.68 12.14 11.96 11.54 11.47 9.31 10.85 11.41 10.39 10.55 10.33 10.17 2.52 2.14 2.23 2.46 2.38 1.49 2.05 2.27 2.97 2.60 2.68 2.59 0.243 0.088 0.078 0.210 0.180 0.475 0.329 0.136 0.125 0.226 0.266 0.624 0.279 0.135 0.154 0.179 0.190 0.203 0.198 0.196 0.213 0.287 0.273 0.399 1.16 0.74 1.77 1.02 0.95 2.63 0.94 1.98 2.68 0.95 2.29 0.94 100.03 100.13 99.96 100.44 99.60 100.14 99.96 100.26 99.96 99.62 100.04 100.16 13.34 10.99 11.03 11.76 12.23 10.37 11.22 11.51 11.82 12.68 12.99 13.33 47.59 59.73 55.92 54.13 51.69 77.91 65.18 58.64 53.23 50.23 47.58 46.65 118 83.4 91.6 99.1 106 76.9 88.5 99.6 108 124 123 125 282 176 171 194 216 126 146 200 178 294 407 270 71.7 132 76.7 84.7 76.6 774 266 220 96.4 110 60.9 68.7 48.7 46.7 36.6 41.5 39.5 27.4 38.2 33.6 35.7 39.0 33.1 34.1 558 422 330 404 403 243 364 350 409 453 354 465 134 381 158 201 173 1049 683 348 277 206 167 92.11 22.6 19.4 19.8 20.8 20.7 13.4 19.2 20.4 20.9 24.0 22.7 25.0 4.50 1.05 1.77 1.53 3.04 10.3 4.79 1.69 1.21 2.95 4.81 10.6 249 222 221 215 224 243 296 228 212 253 274 342 40.9 25.6 26.8 32.3 32.4 20.4 23.1 30.7 32.9 42.5 35.1 39.4 186 103 92.4 123 128 93.5 125 132 140 186 175 232 8.11 4.30 3.55 4.79 5.06 13.5 6.98 5.25 5.43 8.48 11.6 13.5 0.067 0.017 0.109 0.008 0.062 0.035 0.145 0.034 0.028 0.046 0.110 0.124 43.5 19.1 26.4 37.9 33.8 75.6 51.9 39.8 34.9 52.2 93.5 90.0 9.63 5.01 5.05 6.02 6.19 12.52 7.59 6.15 6.52 9.55 12.78 14.54 26.9 14.6 13.5 16.5 17.8 27.3 21.0 17.8 19.7 26.4 32.8 39.4 4.30 2.46 2.08 2.71 2.74 3.44 3.33 2.89 3.04 4.31 4.66 5.88 21.9 12.6 11.0 14.0 14.7 14.9 16.7 15.0 16.5 22.2 22.9 29.5 6.58 3.96 3.56 4.47 4.69 3.66 4.62 4.68 5.15 6.55 6.33 8.08 2.24 1.47 1.32 1.61 1.71 1.20 1.61 1.64 1.83 2.23 2.06 2.62 7.67 4.92 4.31 5.28 5.46 4.19 5.09 5.61 5.96 7.46 7.02 8.66 1.28 0.819 0.747 0.903 0.932 0.635 0.789 0.910 1.04 1.27 1.14 1.33 7.64 4.79 4.53 5.54 5.63 3.66 4.50 5.48 6.10 7.62 6.47 7.61 1.50 0.948 0.932 1.16 1.16 0.755 0.833 1.12 1.22 1.55 1.27 1.42 3.94 2.50 2.53 3.09 3.13 2.03 2.20 2.97 3.24 3.97 3.39 3.74 0.580 0.359 0.367 0.449 0.446 0.280 0.317 0.422 0.470 0.582 0.473 0.508 3.39 2.11 2.30 2.64 2.72 1.76 1.79 2.55 2.81 3.44 2.89 3.02 0.476 0.305 0.320 0.394 0.384 0.252 0.244 0.368 0.399 0.502 0.404 0.415 4.75 2.76 2.51 3.21 3.35 2.36 3.18 3.48 3.73 4.92 4.72 5.76 0.487 0.272 0.250 0.313 0.329 0.801 0.430 0.343 0.371 0.523 0.749 0.886 1.52 0.704 1.17 0.950 0.920 0.731 0.973 1.17 0.918 1.34 2.19 2.19 0.720 0.361 0.424 0.421 0.449 0.868 0.501 0.439 0.412 0.624 1.14 1.01 0.246 0.115 0.110 0.134 0.145 0.401 0.171 0.143 0.117 0.214 0.433 0.337 0.703389 0.703283 0.703318 0.703444 0.703253 0.703354 0.703819 -16.49 -17.99 -17.49 -15.71 -18.41 -16.99 -10.39 0.512945 0.512942 0.512959 0.512794 0.512965 0.512990 0.512785 7.50 7.43 7.78 4.56 7.89 8.37 4.38 17.650 17.719 17.754 18.068 17.982 17.860 17.761 15.346 15.307 15.352 15.533 15.419 15.395 15.301 37.425 37.408 37.552 37.831 37.665 37.661 37.494 lith. code ggu no. deg. w deg. n altitude, m sio2 tio2 al2o3 fe2o3 feo mno mgo cao na2o k2o p2o5 volatiles sum feo* mg number zn cu ni sc v cr ga rb sr y zr nb cs ba la ce pr nd sm eu gd tb dy ho er tm yb lu hf ta pb th u isotope ratios calculated at 60 ma 87sr/86sr60 εsr 143nd/144nd60 εnd 206pb/204pb60 207pb/204pb60 208pb/204pb60 for explanation of lithological codes, see table 1. for petrographical notes on the samples, see table 4c. geographical coordinates in wgs 84. first two digits are degrees, then follow minutes in decimal form. major elements in wt% (xrf analyses). trace elements in ppm (zn–ga: xrf analyses; rb–u: icp-ms analyses). feo* = total iron as feo. mg number = 100 × atomic mg/(mg+fe2+), with the iron oxidation ratio adjusted to fe2o3/feo = 0.15. table 4a. chemical analyses of rocks of the rinks dal member lower rinks dal member skarvefjeld unitunit akuarut unit lower transition flows 111 5131 5131 5131 5131 5141 5141 5151 5161 5171 5171 5181 5181 327049 332861 332852 362227 362168 318758 327023 157261 400255 340908 332945 327014 5324.88 5318.69 5247.61 5228.36 5215.17 5227.90 5326.28 5333.01 5256.76 5216.78 5313.35 5321.72 6916.722 6958.798 7018.247 7012.534 6938.994 7013.426 6916.755 6916.575 7022.688 6938.608 6944.899 6917.239 409 1140 1551 1129 766 1095 620 719 1803 945 1318 829 47.92 47.02 47.86 47.95 48.75 46.43 48.68 48.71 49.33 49.54 48.95 48.17 4.08 4.71 4.51 4.26 2.94 3.00 2.81 1.62 1.98 2.15 2.71 3.87 12.25 12.82 11.85 11.86 12.97 13.75 13.33 14.59 14.51 14.57 13.44 12.22 5.48 6.57 4.94 3.95 3.79 4.97 5.99 5.22 6.00 4.53 4.97 4.82 9.99 8.97 11.33 12.02 10.84 9.10 8.15 6.27 6.11 7.75 9.39 10.90 0.22 0.23 0.26 0.23 0.22 0.22 0.21 0.17 0.18 0.19 0.24 0.23 5.24 5.37 4.66 4.49 5.88 5.46 5.90 7.49 6.72 6.19 6.12 5.58 9.50 9.81 9.16 9.02 10.52 10.69 10.23 12.19 11.60 11.17 10.13 9.57 2.83 2.72 2.71 2.71 2.70 2.79 2.77 2.26 2.43 2.44 2.41 2.67 0.817 0.281 0.688 0.936 0.288 0.962 0.404 0.170 0.153 0.272 0.391 0.472 0.495 0.525 0.620 0.630 0.351 0.410 0.345 0.150 0.184 0.210 0.236 0.458 1.29 1.03 0.96 1.35 0.74 1.83 0.86 0.92 0.65 0.77 1.14 1.22 100.11 100.04 99.54 99.40 99.98 99.61 99.68 99.76 99.84 99.78 100.12 100.18 14.92 14.88 15.78 15.57 14.25 13.57 13.54 10.97 11.51 11.83 13.86 15.24 41.53 42.19 37.40 36.83 45.49 44.86 46.85 58.01 54.15 51.43 47.17 42.55 154 147 164 160 129 115 127 88.0 92.8 101 120 155 354 500 443 396 281 253 304 151 208 218 202 400 55.4 69.5 36.9 36.5 51.7 58.7 53.3 86.0 85.2 63.1 63.6 49.8 36.1 40.8 39.7 30.9 39.1 32.4 34.9 36.3 37.1 39.2 33.7 36.2 506 630 522 416 457 405 419 328 359 404 396 497 62.0 104 39.3 32.4 65.6 65.5 58.6 232 279 171 91.7 75.9 25.9 27.2 26.2 26.2 23.6 21.8 22.9 19.7 20.5 21.0 23.3 24.3 16.5 2.88 16.3 22.6 5.47 19.0 7.65 2.25 0.81 5.71 4.79 5.04 308 347 327 334 263 358 245 212 192 214 220 230 49.5 54.1 56.9 55.1 39.3 35.5 39.9 23.6 30.4 33.8 35.2 55.7 303 348 390 384 186 246 185 82.2 111 128 161 275 20.9 20.1 26.6 26.5 9.11 29.2 8.98 3.36 4.63 6.39 7.84 12.43 0.116 0.064 0.285 0.387 0.084 0.128 0.099 0.010 0.009 0.153 0.053 0.049 130 104 153 155 55.0 190 57.1 24.6 33.0 46.5 74.0 81.0 21.5 21.3 27.4 26.5 9.51 23.5 9.68 3.93 5.32 6.79 9.08 14.4 54.4 57.9 70.5 68.2 27.0 56.0 27.1 11.6 15.0 18.8 24.4 39.5 8.03 9.03 10.78 9.88 4.15 7.45 4.09 1.90 2.55 2.95 3.93 5.82 37.4 43.4 50.4 47.5 21.3 33.8 21.0 10.3 13.3 15.5 19.8 29.9 9.55 11.4 12.5 11.8 6.44 7.95 6.38 3.41 4.38 4.91 5.96 8.88 3.02 3.48 3.69 3.48 2.16 2.44 2.11 1.28 1.56 1.71 2.03 2.81 9.94 12.3 13.7 13.0 7.50 8.35 7.25 4.06 5.10 5.74 6.46 10.3 1.60 1.89 2.05 1.90 1.18 1.20 1.19 0.691 0.885 0.969 1.10 1.65 9.11 10.6 11.4 10.6 7.00 6.65 7.06 4.08 5.45 5.77 6.59 9.80 1.76 1.95 2.10 2.01 1.37 1.29 1.39 0.866 1.14 1.21 1.32 1.94 4.51 5.27 5.39 5.23 3.68 3.43 3.71 2.31 2.99 3.26 3.44 5.33 0.634 0.716 0.765 0.709 0.525 0.463 0.535 0.335 0.441 0.479 0.482 0.751 3.72 4.28 4.50 4.28 3.16 2.86 3.29 1.99 2.58 2.96 2.88 4.58 0.555 0.595 0.632 0.593 0.447 0.407 0.470 0.277 0.395 0.416 0.427 0.663 7.54 8.33 9.52 9.36 4.65 5.82 4.75 2.28 3.02 3.36 4.23 7.03 1.24 1.18 1.57 1.54 0.578 1.67 0.599 0.218 0.336 0.457 0.471 0.769 2.94 2.71 3.49 3.66 1.54 2.46 1.48 0.519 1.04 1.34 1.41 2.50 1.66 1.61 2.38 2.29 0.730 2.23 0.757 0.230 0.403 0.522 0.726 1.19 0.555 0.471 0.777 0.745 0.247 0.634 0.258 0.078 0.127 0.181 0.203 0.415 0.703242 0.703392 0.703182 0.703252 0.703271 0.703390 0.703329 -18.57 -16.45 -19.42 -18.44 -18.16 -16.47 -17.33 0.512962 0.512965 0.512964 0.512993 0.512987 0.512951 0.512978 7.83 7.88 7.87 8.43 8.32 7.61 8.13 17.923 17.957 17.960 17.863 17.987 17.847 17.708 15.404 15.399 15.398 15.383 15.403 15.333 15.338 37.613 37.711 37.711 37.706 37.709 37.591 37.560 lith. code ggu no. deg. w deg. n altitude, m sio2 tio2 al2o3 fe2o3 feo mno mgo cao na2o k2o p2o5 volatiles sum feo* mg number zn cu ni sc v cr ga rb sr y zr nb cs ba la ce pr nd sm eu gd tb dy ho er tm yb lu hf ta pb th u isotope ratios calculated at 60 ma 87sr/86sr60 εsr 143nd/144nd60 εnd 206pb/204pb60 207pb/204pb60 208pb/204pb60 table 4b. chemical analyses of rocks of the rinks dal member upper rinks dal memberakuarut unitunit uppermost flows upper transition flows 112112 354719 very fine-grained basalt with tiny microphenocrysts of plagioclase, clinopyroxene and olivine. foreset-bedded pillow breccia, c. 190 m thick, deposited in the assoq lake. unit 1 in the orpiit qaqqaat profile, south-central disko just east of the gneiss ridge. 332881 very fine-grained, nearly aphyric basalt with sparse microphenocrysts of olivine and plagioclase. from entablature zone in lava flow with well-developed entablature, 25 m thick, resting directly on weathered picrite of the vaigat formation. flow 1 in the kuugannguaq se profile, inner kuugannguaq valley, north-central disko. 326406 fine-grained basalt with numerous plagioclase-olivine glomerocrysts and scattered clinopyroxene phenocrysts. from entablature zone in lava flow with well-developed colonnade and entablature, 30 m thick. flow 4 in the tuapassuit profile, south disko, on depression in the gneiss ridge just west of qeqertarsuaq town. 340961 fine-grained basalt with plagioclase-olivine-clinopyroxene glomerocrysts. lava flow, 60 m thick. flow 4 in the fortunebay east profile, south disko, on depression in the gneiss ridge. 157246 very fine-grained basalt with plagioclase-olivine-clinopyroxene glomerocrysts. from entablature zone in subaqueous lava flow with well-developed colonnade and entablature and thick yellow top breccia, 40 m thick, deposited in the assoq lake. flow 8 in the ippik profile, south disko well east of the gneiss ridge. 136969 picrite with numerous olivine phenocrysts wth inclusions of chromite crystals. pahoehoe lava flow, 7 m thick, one of a group of four picrite flows enriched in some incompatible trace elements. flow 7 in the orlingasoq profile, north disko. 332930 fine-grained basalt with scattered phenocrysts of olivine and clinopyroxene. lava flow, 30 m thick. flow 24 in the kuugannguaq se profile, inner kuugannguaq valley, north-central disko. 362348 fine-grained intersertal basalt with many microphenocrysts of skeletal olivine and a few plagioclases. foreset-bedded pillow breccia, >25 m thick, deposited in the assoq lake. unit 1 in the point 1123 profile, south side of kvandalen, east disko. 136973 fine-grained almost aphyric basalt. lava flow, 5 m thick. flow 2 in the orlingasoq profile, north disko. 362344 fine-grained basalt with sparse plagioclase phenocrysts and olivine microphenocrysts. from entablature zone in 30 m thick lava flow with well-developed colonnade and entablature and red-oxidised top breccia, resting on 10 cm black shale and overlain by 8 m yellow sand (filling stage of the assoq lake). flow 4 in the point 1123 profile, south side of kvandalen, east disko. 327065 well-crystallised basalt with phenocrysts and glomerocrysts of plagioclase. slightly contaminated. lava flow, 12 m thick, with 40 cm thick yellowish brown sediment on top. flow 2 in the skarvefjeld profile, south disko. 279041 very fine-grained basalt with small phenocrysts of plagioclase and olivine and clusters of clinopyroxene. lava flow, 8 m thick. flow 18 in the qullissat profile, north disko. 327049 basalt with scattered plagioclase phenocrysts and glomerocrysts of plagioclase and augite. lava flow, 45 m thick. flow 15 in the skarvefjeld profile, southern disko. 332861 very fine-grained basalt with tiny microphenocrysts of plagioclase, olivine and clinopyroxene. highest tio2 content in the rinks dal member. lava flow, 13 m thick. flow 20 in the kuugannguaq se profile, inner kuugannguaq valley, north-central disko. 332852 very fine-grained basalt with tiny microphenocrysts of plagioclase, olivine and clinopyroxene. lava flow, 14 m thick. flow 5 in the ivissussat qaqqaat profile, east of ataata kuua, south nuussuaq. 362227 very fine-grained basalt with sparse small plagioclase phenocrysts and tiny clinopyroxene microphenocrysts. lava flow, 7 m thick. flow 6 in the giesecke monument profile, south nuussuaq. 362168 fine-grained, nearly aphyric basalt with rare small phenocrysts of plagioclase and clinopyroxene. from entablature zone in 25 m thick lava flow with well developed entablature zone, overlain by c. 2 m of black, coal-bearing sediment. flow 6 in the point 1123 profile, south side of kvandalen, eastern disko. 318758 very fine-grained basalt with scattered phenocrysts of clinopyroxene, plagioclase and olivine, and many microphenocrysts of the same minerals. lava flow, 20 m thick. flow 5 in the giesecke monument profile, south nuussuaq. enriched in some incompatible trace elements. 327023 very fine-grained, nearly aphyric basalt with a few tiny microphenocrysts of plagioclase, olivine and clinopyroxene. lava flow, 45 m thick. flow 21 in the skarvefjeld profile, south disko. 157261 very fine-grained basalt with glomerocrysts of plagioclase and olivine. lava flow, >13 m thick. flow 19 (top flow) in the lyngmarksfjeld profile, south disko. 400255 fine-grained basalt with small plagioclase phenocrysts and plagioclase-olivine-clinopyroxene glomerocrysts. lava flow, 12 m thick. flow 9 in the point 1888 profile, west of ataataa kuua, south nuussuaq. 340908 fine-grained basalt with numerous plagioclase-clinopyroxene-olivine glomerocrysts. lava flow, 55 m thick. flow 9 in the point 1123 profile, south side of kvandalen, east disko. 332945 fine-grained, nearly aphyric basalt with rare plagioclase phenocrysts. lava flow, 25 m thick, with laterite-impregnated top zone, uppermost flow in the rinks dal member. flow 33 in the sorte hak profile, central disko. 327014 very fine-grained, nearly aphyric basalt with scattered small phenocrysts of plagioclase and microphenocrysts of clinopyroxene and olivine. lava flow, 45 m thick, next-highest flow in the rinks dal member. flow 28 in the skarvefjeld profile, south disko. phenocryst phases as observed in thin section are mentioned in order of decreasing abundance. samples of subaerial lava flows are usually taken in massive columns a few metres above the flow base. table 4c. notes on analysed samples of the rinks dal member 113 eruption sites for the rinks dal member very few eruption sites for the rinks dal member are known. it would be natural to assume that the member was fed from the widespread uncontaminated basalt dykes in the region, but this is not the case, as shown below. uncontaminated basalt dykes and sills many dykes of uncontaminated basalt on disko and nuussuaq cut the rinks dal member. in addition, seismic sections and aeromagnetic maps suggest that the sediments beneath disko bugt are intruded by sills; one such sill crops out on a group of small islands in the south-eastern corner of disko bugt (fig. 1), and another has been sampled by dredging in eastern disko bugt (larsen & dalhoff 2007; larsen et al. 2016). south of disko bugt, three large young dykes are known; these are the >60 km long n–s-trending ‘globule dyke’, the 110 km long nne–ssw-trending manermiut dyke and the >13 km long wnw–sse-trending sydostbugt dyke (ellitsgaard-rasmussen 1951; henderson 1969; árting 2004; larsen 2006). on the aeromagnetic maps (rasmussen 2002) the globule dyke and the manermiut dyke form conspicuous linear features with reversed magnetic signature that can be followed northwards into disko bugt and almost to the south coast of disko. no tertiary intrusions have been found on the mainland east of the eastern boundary fault. some dykes are clearly younger than the lava succession because they cut through the whole succession. naturally, such dykes have mainly been identified in the complete lava succession on western disko. such identification cannot be made over large areas, and the age of the dykes must be constrained by other methods. uncontaminated basalt dykes have major element compositions in the same general range as the lavas of the rinks dal member, i.e. 4–8 wt% mgo and 1.4–5 wt% tio2. however, in detail the dykes and sills on disko and nuussuaq are compositionally different from the lavas of the rinks dal member. all dykes with more than 3.2 wt% tio2 (similar to the akuarut unit lavas) have much higher contents of k2o (0.7–1.3 wt%) and p2o5 (0.4–0.7 wt%) and other incompatible elements than the akuarut member lavas and are therefore clearly not feeders for the lava flows. these dykes also have higher ce/y and nb/y than any rinks dal member lavas (fig. 101). most of the cross-cutting dykes on western disko are of this type. they are interpreted to be associated with eocene lavas of the naqerloq formation (larsen et al. 2016). uncontaminated dykes and sills with less than 3.2 wt% tio2 have trace element contents and ratios that are slightly but significantly different from those of the lavas of the rinks dal member. almost all analysed dykes of this kind have lower zr/nb for similar ce/y ratios, and most of them have higher nb/y for similar zr/y ratios (fig. 101) and mainly plot in the iceland field of fitton 0.0 0.5 1.0 1.5 2.0 2.5 0 5 10 15 20 25 30 zrnb rinks dal mb lavas dykes 0.01 0.10 1.00 10.00 1 10 zr/y iceland field ce/y nb/y fig. 101. immobile trace-element ratios of uncontaminated dykes on disko and nuussuaq compared with those of the lavas of the rinks dal member, showing that the dykes were not feeders for the lava flows. three apparent exceptions are discussed in the text. the dykes are interpreted as associated with lavas of the svartenhuk formation (the low-ce/y group) and the naqerloq formation (the high-ce/y group; larsen et al. 2016). 114114 et al. (1998). these dykes are interpreted to be associated with lavas of the latest paleocene svartenhuk formation (larsen et al. 2016). only two intrusions are compositionally similar to the rinks dal member lavas: the manermiut dyke south of disko bugt (two analyses) and an apparent dyke at niuluut on the south coast of disko. the niuluut dyke as observed in 1984 cut a highly fractured subaqueous lava flow with a similar chemistry and was later interpreted as a fracture filling in the flow; the dyke had disappeared completely from the coastal cliff when the locality was revisited in 2011. thus, no analysed dykes on disko and nuussuaq are feeders for the lavas of the rinks dal member. radiometric age determinations confirm the compositional differences because the dated uncontaminated dykes and sills are younger than the maligât formation (fig. 2). the low-ti intrusions are latest paleocene in age, 59–57.5 ma (larsen et al. 2016) whereas the highti intrusions, including the ‘globule dyke’ in the aasiaat district, are earliest eocene in age, 56–54 ma (storey et al. 1998; larsen 2006; larsen et al. 2009). the undated low-ti manermiut dyke (1.7 wt% tio2) in the aasiaat district is considered to be contemporaneous with the rinks dal member lavas, i.e. 61–60 ma. this long, thick dyke, which skirts the outermost west coast south of aasiaat, may be connected northwards to feeder systems for the rinks dal member on western disko and farther west. feeder systems for the rinks dal member only in two localities on western disko, a dyke and a neck have been observed feeding lava flows of the rinks dal member (units 515 and 517, see p. 92–94). as shown above, the dyke systems in the region are unrelated to the rinks dal member, and thus the locations of the feeders for the member largely have to be inferred. several lines of evidence suggest that the main feeder systems are situated west of the disko gneiss ridge: the successive units of the rinks dal member spread eastwards from areas on the gneiss ridge (fig. 15); the lava morphologies are more variable in the west than in the east, the number of flows is higher and the total thicknesses are larger west of the ridge, and the two eruption sites that have been found are both situated west of the ridge. the lack of potential feeder dykes east of the gneiss ridge points in the same direction. west of the gneiss ridge, the feeders will be covered by younger flows, but east of the ridge, erosion has cut sufficiently deep to expose them, if they were present. already henderson (1973) speculated that the primary source areas of these basalts were on western disko and west of disko. the skarvefjeld unit (unit 511) may have been erupted from more easterly sites. the very thin pahoehoe flows of this unit that filled the assoq lake on eastern disko could not have flowed for very long distances and may be fed from sites east of the gneiss ridge. support for this may be found in the distribution of geochemically enriched flows of unit 511: these occur in a limited area on eastern north disko around orlingasoq, where the strongly enriched flows of the manîtdlat member of the vaigat formation also occur (larsen et al. 2003; pedersen et al. 2017). the enriched unit 511 magmas could have acquired their composition during passage through old feeder systems for the manîtdlat member in the crust, leaving the main, deep crustal magma chambers unaffected. another small succession of thin pahoehoe flows (unit 507; flows βfph1 on the map of pedersen et al. 2001) only known from a 20 × 6 km area in the kuannersuit kuussuat valley on central disko was probably also erupted locally. 115 nordfjord and niaqussat members revision of the boundary between the nordfjord and niaqussat members the upper part of the maligât formation is lithologically very variable and comprises crustally contaminated rocks ranging from basalts and picrites to andesites, dacites and rhyolites. the rocks form lava flows, high-level intrusions, crater breccias and rhyolitic tuff layers, with crater areas centred on western disko. pedersen (1975a, 1977a, b) divided this succession into two members, the nordfjord and niaqussat members, based on lithological differences between the basic rocks coupled with compositional differences. the nordfjord member was restricted to comprise relatively evolved basalts together with silicic rocks derived from evolved basaltic parent magmas, whereas the niaqussat member also comprised picrites and magnesian basalts, as well as more silicic rocks with magnesia-rich parents. included in the niaqussat member were silicic rocks in a number of craters and feeder dykes on north-western disko, with niaqussat as the type area for the member. the niaqussat member was interpreted to represent a major new influx of primitive picritic melt into the main crustal magma reservoir. subsequent investigations have shown that the original distinction between the two members is difficult to uphold, particularly for the more silicic rocks. pedersen (1981) showed that some of the andesites and dacites of the niaqussat member have evolved basaltic parents. moreover, at various stratigraphic levels, the nordfjord member comprises subordinate magnesian basalts with up to 10.5 wt% mgo; this indicates that the influx of primitive magma that culminated with the basic magmas of the niaqussat member with up to 15 wt% mgo had already begun in nordfjord member time. in consequence, we here revise the boundary between the nordfjord and niaqussat members as follows. the base of the niaqussat member is now placed at the base of a succession of flow-folded pahoehoe lavas of olivine-microphyric magnesian basalts and picrites. this succession is lithologically distinctive and has a large regional extent over all of disko and parts of eastern nuussuaq. all the crater facies rocks on north-western disko, including the niaqussat crater, as well as a large composite native-iron-bearing lava flow in the mellemfjord area, are thereby moved to the nordfjord member. this leaves no andesites and dacites, and only few basaltic andesites, in the niaqussat member. the revision has no consequence for the boundary between the two members on central and eastern disko and nuussuaq, which stays in place. it also has no consequence for the geological map sheets on a scale of 1:100 000 because these combine the nordfjord and niaqussat members on western disko. the two members are shown separately on the geological map of the mellemfjord area on a scale of 1:50 000 (pedersen 1977b) and in the photogrammetric sections on a scale of 1:20 000 (south disko section, central disko section), and the only consequence for these is that the large native-iron-bearing mellemfjord lava flow, which appears as the lowest flow of the niaqussat member on the 1:50 000 scale map and in the south disko section at 8–16.3 km, is now considered to be the uppermost flow of the nordfjord member (fig. 102). 200 11 12 13 14 15 km 400 600 800 1000 m point 639 m inngissuaq fe fe fe d d532 530 531? 520 fe cr gv03_03_060_lml 530 531? 520 wnw ese fig. 102. photogrammetrically measured section of the coastal cliff at eqaluit on south-west disko, documenting the boundaries between the rinks dal, nordfjord and niaqussat members. after the present revision of the boundary between the nordfjord (520) and niaqussat (530–532) members, the native-iron-bearing lava flow (fe) is shifted from the lowest niaqussat member to the uppermost nordfjord member (this flow is the composite mellemfjord lava flow described later). green flows beneath the nordfjord member belong to the upper rinks dal member. some characteristic marker flows are separately coloured. an eruption crater (cr) in unit 532 is traced with a white outline. two dykes (d) cut the succession, one very obliquely. slightly modified excerpt from the south disko section (pedersen et al. 2003). 116116 nordfjord member summary of the main features of the nordfjord member • original extent over all of disko and southern and eastern nuussuaq, with depocentre on western disko. • cannot be subdivided based on stratigraphy. descriptions based on rock types and localities. • very wide igneous compositional range, from basalt to rhyolite, with no stratigraphic significance. numerous tuff layers. • all igneous rocks are more or less crustally contaminated except for one alkali basalt flow. sediment xenoliths are common in the more silicic lavas and tuffs. • the dominant lithology is basalt lava flows. andesites, dacites and rhyolites are confined to north-west and central-west disko where eruption sites are known and inferred. • native iron occurs in some basaltic andesites, andesites and dacites, and rarely in basalts. • composite lavas with more basic lower parts and more silicic upper parts, and many with native iron, occur on west disko and northern east disko. they indicate the presence of several individual high-level magma chambers within sediments west of the disko gneiss ridge and, on northern east disko also east of the ridge. • an unexposed central volcano, the west disko graphite rhyolite volcano, is situated somewhere on northwest disko or just offshore from there. the volcano produced massive rhyolites (now only found in conglomerates) and graphite-bearing tuffs which are the most strongly sediment-contaminated and evolved rocks of the maligât formation. • a metre-thick sediment deposit at the base of the member suggests an interruption of the volcanic activity after the formation of the rinks dal member. • volcanogenic conglomerates and sandstones with plant fossils occur at several levels on north-west disko and indicate an active erosive environment with river transport. picrite blocks in the conglomerates indicate exposure and erosion of the vaigat formation to the north. lithostratigraphy of the nordfjord member revised member history. the nordfjord member was informally established by pedersen (1975a) in the area between hammer dal and nordfjord on north-western disko. it was further investigated on north-western disko by pedersen (1977a) and on western and south-western disko by pedersen (1977b, 1981) who gave the first detailed descriptions. the nordfjord member was subsequently mapped on southern, south-eastern and eastern disko (pedersen & larsen 1987; larsen & pedersen 1989) and on nuussuaq (pedersen & larsen 1987; larsen & pedersen 1992). the definition is here formalised and extended to cover the whole of disko and nuussuaq. name. after nordfjord, north-western disko (fig. 4). distribution. the nordfjord member originally extended over the whole of disko and southern and eastern nuussuaq (fig. 15); its main occurrence now is in the downfaulted, west-dipping blocks of western disko where it is preserved over wide areas (fig. 103). it is eroded away on central and north-central disko (fig. 4), whereas on eastern and southern disko it is present as erosional remnants on peaks and ridges, particularly on nunataks in and around the central glacier sermersuaq (e.g. pedersen et al. 2001, 2003, 2005). it caps the highest peaks on southern nuussuaq (point 2010 m and giesecke monument) and is present at high altitudes on eastern nuussuaq (larsen & pedersen 1992). type section. point 440 m, northern gully, on northwestern disko (fig. 14, profile 11). the locality is an e–w-oriented gully 3.5–4 km north of the river in hammer dal (fig. 103); the succession is faulted and tilted and dips 22°w (photogrammetrically measured section in fig. 104a). the nordfjord member is exposed over a lateral distance of c. 2000 m at around 350 m altitude. reference sections. point 440 m, southern gully, an e–wrunning gully c. 1.2 km south of the gully with the type section (fig. 14, profile 10; figs 103, 104b). niaqussat at 280–380 m altitude, north-western disko (fig. 14, profile 12). point 600 m, north side of hammer dal (fig. 117 14, profile 9; figs 103, 104c). sedimentkløften at the south side of the mouth of hammer dal, north-western disko (fig. 117). point 1070 m at c. 950–1200 m altitude, north of outer nordfjord, western disko (fig. 14, profile 8; central disko section at 7.5–9 km). sapernuvik at c. 50–140 m altitude, west coast of disko north of kangerluk (fig. 14, profile 1; south disko section at 4 km). the eastern side of the mountain akuliarusersuaq at 1045–1080 m altitude, central disko (fig. 9, profile 2). the eastern wall of the kuugannguaq valley just south of 70°n, at 1600–1670 m altitude, northern disko (fig. 9, profile 5; central disko section at 55–57 km). the southern shoulder of the mountain qinngusaq above the innermost part of kvandalen, at c. 1150–1290 m altitude, eastern disko (fig. 10, profile 8; central disko section at 81–83 km). the mountain point 1888 m west of ataata kuua, at 1860–1893 m altitude, 7.5 km from the south coast of nuussuaq (fig. 11, profile 2). the southern slopes (nunavik), at c. 1650–1680 m altitude, of the mountain point 2000 m, eastern nuussuaq (fig. 12, profile 13; central nuussuaq section at 67–68 km). saqqaq, point 1266 m at 1235–1270 m altitude, 10 km north of the village saqqaq, south-eastern nuussuaq (fig. 11, profile 11). thickness. the nordfjord member has its depocentre on western disko where it is up to 290 m thick and where 10 or more lava flows are present in some profiles. another, smaller, depocentre seems to exist on eastern disko around kvandalen where the member is up to 180 m thick and comprises nine lava flows in the qinngusaq section. in most other places on disko where the upper boundary to the niaqussat member is preserved, the nordfjord member is less than100 m thick and comprises only 2–5 lava flows. in many profiles only 1–2 flows are left after erosion. on eastern nuussuaq, the nordfjord member is only about 30 m thick and comprises 2–3 lava flows. lithology. lava flows comprise plagioclase-phyric tholeiitic basalts, plagioclase-phyric to aphyric basaltic andesites and andesites some of which carry native iron, and weakly plagioclase-orthopyroxene-phyric dacites, some with native iron. some lava flows are composite. one flow is an alkali basalt. quartz-feldspar-biotite-phyric rhyolite, some with almandine garnet, forms tuff layers. all lava flows are subaerial. volcaniclastic sediments occur at several horizons and comprise conglomerates with clasts of dacite and rhyolite, and sandstones of redeposited tuffaceous material, some with plant remains. sediment xenoliths showing variable degrees of reaction with the magma are common in the silicic lavas and tuffs. subdivisions. the nordfjord member has not been subdivided. boundaries. the nordfjord member conformably overlies the lava flows of the rinks dal member; as described below there is commonly, but not always, a sediment horizon at the base. at the upper boundary, the nordfjord member is conformably overlain by a distinctive succession of flow-folded pahoehoe lavas of olivine-microphyric magnesian basalts and picrites of the niaqussat member. age. paleocene, 61–60 ma, magnetochron c26r, based on radiometric dating (storey et al. 1998; larsen et al. 2016). correlation. none certain. internal structure of the nordfjord member despite its modest thickness, the nordfjord member is the most variable and complex member in the maligât formation. in contrast to the rinks dal and niaqussat members with their laterally continuous units, the nordfjord member contains a patchwork of local eruption sites surrounded by various lithologies of limited areal extent. a subdivision into chemostratigraphic units is consequently not possible, and the entire member is referred to as unit 520. the description of the geology of the nordfjord member in the following is therefore structured differently from the descriptions of the other members. the nordfjord member descriptions comprise, firstly, a section with descriptions of the different igneous rock types and, secondly, a section arranged according to geological themes illustrated by many locality descriptions. many of these localities are situated on north-western disko and their locations are shown in fig. 103. 118118 5 4 7 5 5 4 2 3 4 4 3 2 3 2 2 4 4 4 3 3 5 11 1014 5 7 7 7 77 1020 37 7 5 4 1070 m 26 3 3 3 3 2 42 2 4 7 15 16 21 1333 20 5 15 8 16 13 2232 37 25 30 27 27 34 14 12 20 22 22440 m 500 m 18 10 8 7 4 15 32 ? ? ? gv01_02_158_lml 5 km niaqussat 1 2 3 4 5 a b c 66 8 9 10 11 12 13 14 27 16 18 17 19 20 21 22 23 24 25 15 7 jam m a nordfjord and niaqussat mbs contaminated dykes and craters late dolerite intrusions quaternary 3 locality rinks dal member vaigat formation nordf jord st or da l st ee ns tr up d al g ie se c k e dal rink dal hammer dal 1460 m 1560 m 1510 m 1590 m 1610 m 1380 m hanekammen 1132 m 28 30 29 31 32 33 34 rink dal 882 m 70°15'n 54°40'54°50' 54°20'54°30' 54°10' 70°10'n 70°05'n 70° 119 igneous rock types the major part of the nordfjord member consists of subaerial basalt lava flows. lava flows and crater deposits of basaltic andesite, andesite and dacite, some of which carry native iron, are subordinate and mainly occur on western disko except for a few flows of basaltic andesite on eastern disko. some lava flows are composite, typically with more basic lower parts and more silicic, commonly native-iron-bearing main parts. the most characteristic component of the nordfjord member, although volumetrically insignificant, is rhyolite, which solely occurs as tuffs and conglomerate blocks on western disko and at one locality on north-eastern disko. basalt composition and petrography. the basalt lava flows of the nordfjord member are commonly very similar to the basalts of the underlying rinks dal member. compositionally, the nordfjord member lavas are ordinary tholeiitic basalts; compared to the rinks dal member basalts they show small but systematic differences (see below) suggesting that almost all the nordfjord member basalts are slightly crustally contaminated, as discussed by larsen & pedersen (2009). they show wide ranges in mgo (5.3– 10.0 wt%) and tio2 (1.3–3.9 wt%). basalts with 50–52 wt% sio2 are also called silicic basalts. petrographically the basalts vary from aphyric to strongly plagioclase-phyric. the most magnesian basalts (fig. 105a) are olivine-microphyric and may in addition carry scarce microphenocrysts of plagioclase. there are many almost aphyric basalts which comprise a considerable compositional range. several basalts carry microphenocrysts of plagioclase, augite and scarce olivine and are petrographically indistinguishable from typical rinks dal member basalts. the lithologically most distinctive nordfjord member basalts, which are widespread and common, are highly plagioclase-phyric with many individual phenocrysts larger than 5 mm and glomerocrysts up to 15–20 mm in size (fig. 105b). these basalts also contain phenocrysts of augite and olivine and represent a range in tio2 of 2.2–3.0 wt%. the silicic basalts may also carry microphenocrysts of low-ca clinopyroxene or orthopyroxene. the slight to strong crustal contmination of the nordfjord member basalts is also demonstrated by the widespread occurrence of xenocrysts and xenoliths of crustal (mostly sedimentary) origin. some xenoliths comprise plagioclase-spinel-graphite aggregates derived from mudstone. facing page: fig. 103. geological map of north-west disko showing mentioned localities and figured photographs. localities 1–26 are numbered in a general direction from north to south; localities 27–34 are scattered. 1: logs c and d in fig. 113, photos in figs 114–116. 2: airfall tuff locality 3 km north-north-east of point 440 m north of hammer dal; log in fig. 117, photos in fig. 122. 3: photo in fig. 129. 4: purple line: ‘point 440 m, northern gully’, type section for the nordfjord and lower and middle niaqussat members, fig. 104a. 5: figures from the type section: base of the nordfjord member log b in fig. 113, airfall tuff log in fig. 117, photos in figs 121, 123b, 127. 6: purple line: ‘point 440 m, southern gully’, reference section for the nordfjord and niaqussat members, fig. 104b. 7: point 440 crater, photos in figs 143, 144. 8: point 600 m, photo in fig. 132. 9: purple line: ‘point 600 m, north side of hammer dal’, reference section for the nordfjord member, fig. 104c. 10: hammer dal complex, drawing and photo in figs 175, 176. 11: hammer dal transgressive sill, drawing and photos in figs 177–179. 12: sedimentkløften with tuffaceous sediments, log in fig. 117, photos in figs 118–120. 13: conglomerate, photo in fig. 128. 14: volcaniclastic sandstone, photo in fig. 131. 15: prominent conglomerate, photo in fig. 130. 16: picritic lava succession of the niaqussat member, photo in fig. 151. 17: hanekammen complex, neck, photo in fig. 180, map in fig. 181. 18: niaqussat member basalt lava flow packed with xenoliths, photo in fig. 154. 19: hanekammen complex, crater, map in fig. 181, photo in fig. 182, drawing in fig. 183. 20: nordfjord complex, map in fig. 181. 21: nordfjord complex, tubular intrusion, photo in fig. 184. 22: the point 1070 section (fig. 14, profile 8) with sediments at the base of the nordfjord member, log a in fig. 113. 23: nordfjord member lava flows, photo in fig. 134. 24: nordfjord complex, important sediment xenolith locality. 25: rinks dal member, volcanic neck in unit 517, photo in fig. 84. 26: the kingittuusaq section (fig. 14, profile 7). 27: tuff locality, 2.5 km ssw of point 521 m (avatarpaat qaqqaat). 28–29: inner giesecke dal dyke zone. 30: outer giesecke dal dyke. 31–34: innermost hammer dal dyke swarm. 20 0 fe fe fe fe fe c gl. c gl. t t t t t fe 17 0° /2 2° w c ra te r b as e o f gu ll y b as e o f gu ll y 60 0 80 0 40 0 10 00 12 00 14 00 w 50 0 m 40 0 30 0 e 20 0 m po in t 4 40 m , n or th er n gu lly ty pe se ct io n fo r t he n or df jo rd m em be r a nd th e lo w er a nd m id dl e n iaq us sa t m em be r g v0 3_ 03 _0 48 _l m l 14 00 18 00 16 00 22 00 20 00 24 00 26 00 w e 50 0 m 40 0 30 0 16 2° /2 2° w a a fi g 12 7 fi g 12 1 fi g 11 3 fi g 11 7 ? ? ? ? ? ? 121 20 0 m 20 0 m g v0 3_ 03 _0 49 _l m l 0 40 0 60 0 20 0 80 0 10 00 40 0 60 0 fe fe fe 20 0 80 0 w e n iaq us sa t m em be r so lif lu ct io n an d sc re e ba sa lt lav a flo w s, un it 53 1 ba sa lt an d pi cr ite la va fl ow s, un it 53 0 m as siv e gr ap hi te -r ich d ac ite la va fl ow m as siv e da cit e lav a flo w an de sit e lav a flo w s an de sit e cr at er sc or ia an de sit e tu ff ba sa lti c an de sit e lav a flo w s sil ici c ba sa lt lav a flo w s c on glo m er at e ric h in rh yo lit e bl oc ks g ra ph ite -b ea rin g rh yo lit e tu ffs la te rit ic so il ba sa lt lav a flo w s, un it 51 8 ba sa lt lav a flo w s, un it 51 7 ba sa lt lav a flo w s, un di ffe re nt iat ed d ol er ite in tr us io n te ct on ic br ec cia g ra ph ite -r ich d ac ite c ra te r b re cc ia 50 0 m 40 0 30 0 w e 60 0 m 50 0 40 0 fe c gl. t fe fe fe fe ? 24 °w c ra te r c ra te r fe fe fe po in t 4 40 m , s ou th er n gu lly po in t 6 00 m , n or th si de o f h am m er d al n at ive ir on fe fa ul t n or df jo rd m em be r ri nk s d al m em be r b as e o f gu ll y b c di p ? fi g. 1 04 (p ag es 12 0, 1 21 ). ph ot og ra m m et ric d ra w in gs o f t ilt ed an d fa ul te d se ct io ns n or th o f h am m er d al , w es te rn d isk o. a : p oi nt 4 40 m , n or th er n gu lly (s ho w n in tw o ov er la pp in g p an el s), th e t yp e s ec tio n fo r th e n or df jo rd m em be r a nd th e l ow er an d m id dl e n ia qu ss at m em be r. b : p oi nt 4 40 m , s ou th er n gu lly . c : p oi nt 6 00 m . b an d c ar e r ef er en ce se ct io ns . c om pa re th e r ec on st ru ct ed p ro fil es in f ig . 1 4, p ro fil es 9 –1 1, an d th e p ho to gr ap h of p oi nt 6 00 m in f ig . 1 32 . t he tr ac es o f t he th re e p ro je ct io n pl an es ar e s ho w n in f ig . 1 03 . 122122 over a distance of around 5 km along the north coast of nordfjord, the first lava flow above the basal sediment horizon is an alkali basalt, the only alkaline rock in the entire maligât formation. pebbles of this flow occur in a conglomerate in the hammer dal area c. 7 km north of the exposures of the flow. the alkali basalt is a very finegrained, aphyric rock and differs from other basalts by containing traces of dark mica in micropegmatitic vugs in slowly cooled parts of the lava. in a few profiles on western disko, notably in the giesecke dal and niaqussat profiles (figs 13, 14), the first lava flow above the basal sediment horizon is a highti basalt flow with 3.3–3.9 wt% tio2, similar to some of the flows in the uppermost unit (518) of the rinks dal member. however, the sediment horizon provides an unequivocal boundary between the two members. distribution. basalt lava flows constitute the volumetrically dominant rock type in the nordfjord member and are present throughout the member on disko and nuussuaq (fig. 15; see also e.g. figs 155, 156 and 159). basaltic andesite composition and petrography. by definition, the basaltic andesites have 52–57 wt% sio2; the transition from basalt is gradual. they show wide ranges in mgo (3.4–10.6 wt%) and tio2 (1.2–3.2 wt%). the majority of the basaltic andesites carry phenocrysts of plagioclase and orthopyroxene and frequently xenocrysts of plagioclase and plagioclase-spinel aggregates. several of these rocks carry native iron and sulfides, and many have a very fine-grained groundmass with a globular texture (fig. 106a). patches of decomposed magmaequilibrated mudstone are common in these rocks. some of the most magnesian rocks of the nordfjord member are basaltic andesites in vesterdalen/qasigissat kuussuat. these rocks have more than 800 ppm cr and are clearly derived from a much more primitive magma than represented by the basalts of the upper rinks dal member. the glassy sample 263934 (fig. 106b) is one of the most magnesian basaltic andesites and the only one mf098a_8cm327011polmodcut.eps a pl 5 mm ol mf098b_8cm332964upmodcut.eps b pl ze cpx 5 mm ol fig. 105. thin sections (scanned) of basalts of the nordfjord member. a: basalt with olivine microphenocrysts (ol, pseudomorphed) and a plagioclase phenocryst (pl) in a fine-grained groundmass. sample ggu 327011, skarvefjeld, south disko; sio2 = 50.2 wt%, tio2 = 1.69 wt%, mgo = 8.9 wt%. b: characteristic nordfjord member basalt type with many large phenocrysts and glomerocrysts of plagioclase (pl), augite (cpx) and smaller olivine crystals (ol, pseudomorphed) in a fine-grained groundmass. rounded vesicles are filled with zeolite (ze). sample ggu 332964, nunatak near point 1440 m, central disko; sio2 = 49.2 wt%, tio2 = 2.90 wt%, mgo = 6.0 wt%. concentrations are recalculated volatile-free. 123 with fresh, well-preserved phenocrysts and microphenocrysts of olivine. in addition, the rock contains scarce microphenocrysts of orthopyroxene, spongy plagioclase xenocrysts and tiny microphenocrysts of clinopyroxene. the glass contains 55.4 wt% sio2 and 6.5 wt% mgo. one of the least magnesian basaltic andesites is exemplified by sample 332944 with 53.3 wt% sio2 and 5.8 wt% mgo. this rock is very fine-grained, flow-laminated, and virtually aphyric except for very scarce plagioclase microphenocrysts (not illustrated). distribution. basaltic andesites mainly occur on western and north-western disko. here they form simple lava flows intercalated with basalt or andesite flows, as in the hammer dal area (fig. 14), or several successive flows, as in vesterdalen. basaltic andesite also forms part of composite lava flows, as described later. a few basaltic andesite flows occur on southern and eastern disko and nuussuaq. at uiffaq, tunup qaqqaa, daugaard-jensen dal, sorte hak and point 1888 m, the first lava flow in the nordfjord member is a basaltic andesite (figs 7–9, 11). in the kvandalen area on easternmost disko, an up to 30 m thick flow of basaltic andesite with native iron occurs in the middle of the nordfjord member (fig. 10). andesite composition and petrography. by definition, the andesites have 57–63 wt% sio2; the transition from basaltic andesite is gradual. they show a wide range in mgo (2.4– 7.9 wt%) and less in tio2 (1.1–2.3 wt%). the andesites range from very fine-grained, nearly aphyric rocks to very phenocryst-, xenocrystand antecryst-rich, native-iron-bearing rocks exemplified by the andesite erupted from a crater at niaqussat (fig. 107a). the phenocrysts include orthopyroxene, pigeonite, plagioclase and armalcolite. there are traces of ilmenite, rutile, native iron, troilite and graphite, and xenocrysts of plagioclase, spinel, corundum, cordierite, mullite, quartz and graphite. there is an abundance of magma-modified xenoliths of mudstone and sandstone, and cognate clusters of relatively coarse noritic rocks. mf099a_8cm264076dglmofcut.eps a pl fe xen opx 5 mm mf099b_8cm263934polmodcut.eps b 5 mm fig. 106. thin sections (scanned) of basaltic andesites of the nordfjord member. a: native-iron-bearing magnesian basaltic andesite with phenocrysts of plagioclase (pl) and microphenocrysts of orthopyroxene (opx) in a very fine-grained groundmass with globular texture and blebs of native iron (fe). a xenolith (xen) is seen in the lower right corner. sample ggu 264076, vesterdalen, west disko; sio2 = 56.1 wt%, tio2 = 1.53 wt%, mgo = 6.8 wt%. b: magnesian basaltic andesite, nearly aphyric glassy rock with brown alteration rims of the glass along cracks and vesicle walls. sample ggu 263934, pahoehoe tongue, point 1114 m, vesterdalen, west disko; sio2 = 55.2 wt%, tio2 = 1.58 wt%, mgo = 9.2 wt%. concentrations are recalculated volatile-free. 124124 a low-al2o3, high-p2o5 variety of andesite associated with similar dacite (see below) contains microphenocrysts of tridymite in a glassy groundmass (fig. 107b). distribution. andesites in the nordfjord member are restricted in occurrence to western and north-western disko from mellemfjord to niaqussat (fig. 108). the niaqussat crater erupted a thick andesite flow (fig. 14, profile 12). dacite composition and petrography. the dacites have sio2 in the range 63–67.7 wt% and contain 1.3–3.0 wt% mgo and 1.3–2.3 wt% tio2. there are four main types of dacite, which are distinguishable both chemically and petrographically. native-iron-bearing dacite with ilmenite. this is the most voluminous dacite type. it is confined to north-western disko around and north of hammer dal, where it occurs as a few lava flows and as pebbles in conglomerate. three samples of this type were described in detail by pedersen (1981). the type is exemplified here by sample 176486 (fig. 109a); this is a very fine-grained rock with phenocrysts of plagioclase and orthopyroxene and scarce microphenocrystic ilmenite, which has been modified by progressive reduction and sulfidation to aggregates of ilmenite, armalcolite, rutile, native iron and troilite (fig. 110). there are small, dispersed grains of native iron and troilite and flakes of graphite in the groundmass. xenoliths and xenocrysts derived from mudstone and sandstone and antecrysts and cognate noritic inclusions are widespread. mf100a_8cm176411xdglmodcut.eps a pl pl opx opx opx opx pl 5 mm mf100b_8cm263910polmodcut_c.eps b glass opx pl pl opx 5 mm fig. 107. thin sections (scanned) of andesites of the nordfjord member. a: andesite (lower and upper part of the image) with phenocrysts of plagioclase (pl) and orthopyroxene (opx) in a fine-grained groundmass with tiny black specks of native iron. most of the image is occupied by a cognate inclusion of orthopyroxene and plagioclase with a dark brown, altered residuum. detailed description and mode of this sample in pedersen (1981). sample ggu 176411, niaqussat, north-west disko; sio2 = 60.6 wt%, tio2 = 2.02 wt%, mgo = 2.51 wt%. b: low-al2o3, high-p2o5 andesite with microphenocrysts of orthopyroxene and xenocrysts of sieve-textured plagioclase in a glassy groundmass. pie-sector-shaped microliths of tridymite in the groundmass are too small to be discernible in this image. sample 263910, qasigissat, west disko; analysed bulk rock 263911 has sio2 = 60.7 wt%, tio2 = 2.00 wt%, mgo = 4.45 wt%. 125 a closely related native-iron-bearing dacite with 66– 67 wt% sio2 is only known from a coastal exposure at jamma. it is the most silicic igneous rock with native iron known from the nuussuaq basin and has been described in detail by pedersen (1981). it carries phenocrysts of orthopyroxene, pigeonite, plagioclase and ilmenite and a range of xenocrysts derived from sediments and sediment xenoliths, and cognate noritic inclusions (fig. 109b). low-al2o3 , high-p2o5 dacite. dacite of this type is characterised by relatively high p2o5 and tio2 and low al2o3 and k2o. it is confined to western disko between mellemfjord and the north coast of nordfjord (fig. 108c), where it forms part of a composite lava flow with a range of compositions from high-si andesite to dacite that share similar chemical characteristics. the dacitic rocks are generally very fine-grained and phenocrystpoor but may contain up to a few per cent phenocrysts of orthopyroxene and plagioclase and scarce microphenocrysts of clinopyroxene and pseudomorphed olivine. the orthopyroxene phenocrysts commonly have reaction rims of clinopyroxene. resorbed xenocrysts of quartz and quartz aggregates and scattered, sieve-textured plagioclase grains are common, whereas fine-grained plagioclase-spinel-graphite microxenoliths are scarce. most rocks contain scarce native iron and troilite. the groundmass, especially when well crystallised, reveals the presence of tridymite in addition to plagioclase, clinopyroxene, ilmenite and residuum, and in glassy rocks, the tridymite is seen to be an early crystallising phase. this magma type is inferred to have been formed by the reaction of evolved basalt (delivering the p2o5 and tio2) with a sandstone-dominated contaminant (delivering the quartz xenocrysts). high-ni dacite. dacite of this type is characterised by relatively low tio2 and cao and significantly higher ni than in the other dacites (c. 200 ppm, see table 6 (pages 164–168), sample 176443). it forms a thick crater deposit with an associated lava flow in the hammer dal area where it is the highest unit in the nordfjord member (fig. 14, profiles 9, 10). the dacite has phenocrysts, glomerocrysts and microphenocrysts of orthopyroxene and plagioclase together with minor ilmenite, armalconordfjord st or da l gv01_02_159_lml 25 km malig aat mellemfjord 70°n 55°w nordfjord st or da l malig aat mellemfjord 70°n 55°w nordfjord st or da l mellemfjord 70°n 55°w nordfjord st or da l malig aat mellemfjord 70°n 55°w composite mellemfjord flow, basalt-to-andesite with native iron. two composite basalt-to-basaltic andesite flows with native iron. red contour: composite andesiteto-dacite flows with low al and high p. green contours: andesite flows with native iron. dacite flows with native iron, ilmenite and armalcolite. niaqussat jamma a b c d fig. 108. distribution of some composite flows, andesite flows and dacite flows in the nordfjord member, outlined in red (a–d) and green (c); dashed lines are inferred boundaries. blue dots are the location of profiles shown in fig. 14, but several samples from outside these profiles were also used to construct the figure. there are no systematic age differences between the flows. 126126 mf102a_8cm176486rpolcut_c.eps a 5 mm ilm xen mu opx pl pl pl+opx mf102b_8cm176471xdglmodcut_c.eps b 5 mm pl xpl ilm pl opx opx mf102c_8cm176443rpomod2_c.eps c 5 mm xen opx opx opx+pl pl pl opx pl+sp+gr pl+sp+gr mf102d_8cm326550polmodcut_c.eps d 5 mm sp pl+gr +sp pl+gr +sp opx+pl pl pl opx fe-ti ox mul+gr 127 lite and rutile (fig. 109c). it also contains native iron, troilite and disseminated graphite. the very fine-grained groundmass contains disseminated flakes of graphite and has a composition transitional between dacite and rhyolite, with 68–71 wt% sio2. there is an abundance of mediumto fine-grained igneous clasts of picrite and noritic dolerite together with a large range of magma-modified xenocrysts and xenoliths derived from mudstone and sandstone. xenocrysts of resorbed quartz and cordierite rimmed by plagioclase are common, together with plagioclase-spinel-cordierite-mullite-graphite rocks and many other similar rocks. scarce plagioclase-quartz-pyroxene rocks with glassy residuum and tiny flakes of phlogopitic mica and graphite also occur. residual glass within the xenoliths is rhyolitic. high-al graphite dacite. dacite of this type is characterised by relatively high al2o3 and mgo and low p2o5 and feo*. graphite is abundant. it is only found as cobbles in a nordfjord member conglomerate just north of hammer dal. it has not been located in situ and is speculated to have originated from the west disko rhyolite volcano. it is a very fine-grained volcanic rock with phenocrysts of orthopyroxene and plagioclase in a dark groundmass with about 2.7 wt% toc (total organic carbon) in dispersed graphite flakes. there are abundant small xenocrystic aggregates of aluminous phases and graphite in various stages of equilibration with the magma. the aluminous phases include mullite, cordierite, aluminous spinel and corundum and are derived from carbonaceous mudstone. rare xenocrystic aggregates with calcic clinopyroxene and without graphite are probably derived from sandstone. the rock is inhomogeneous with a veined facing page: fig. 109. thin sections (scanned) of dacites of the nordfjord member. a: dacite with phenocrysts of plagioclase (pl), orthopyroxene (opx) and scarce microphenocrystic ilmenite (ilm). the groundmass contains small dispersed grains of native iron and troilite (both weathered to brown rusty spots) and flakes of graphite (tiny black specks). a re-equilibrated xenolith (xen) consists of mullite (mu) rimmed by plagioclase (pl). detailed description and mode of this sample in pedersen (1981). sample 176486, clast in conglomerate north of hammer dal, north-west disko; sio2 = 65.4 wt%, tio2 = 1.52 wt%, mgo = 1.42 wt%. b: dacite with phenocrysts of plagioclase (pl) and orthopyroxene (opx) as well as xenocrysts (x) derived from sediments and sediment xenoliths. native iron forms tiny black specks in the groundmass. a cognate noritic inclusion is seen at the top. detailed description and mode in pedersen (1981). sample 176471, jamma, north-west disko; sio2 = 67.7 wt%, tio2 = 1.28 wt%, mgo = 1.31 wt%. c: high-ni-dacite with phenocrysts, glomerocrysts and microphenocrysts of orthopyroxene (opx) and plagioclase (pl). native iron forms small black specks in the groundmass. several xenocrystic clusters of plagioclase, spinel and graphite (pl+sp+gr) are seen. a picrite xenolith (xen) is present near the base. sample176443, hammer dal, north-west disko; sio2 = 66.0 wt%, tio2 = 1.29 wt%, mgo = 2.11 wt%. the sample contains 1.14 wt% carbon. d: high-al graphite dacite with phenocrysts of orthopyroxene (opx) and plagioclase (pl) in a dark groundmass with dispersed graphite flakes. small xenocrystic aggregates consist of aluminous phases and graphite (gr). the aluminous phases include mullite (mul), cordierite and red aluminous spinel (sp). the rock is inhomogeneous due to the mixing and mingling of more and less contaminated magma. sample 326550, clast in conglomerate north of hammer dal, north-west disko; sio2 = 63.5 wt%, tio2 = 1.76 wt%, mgo = 3.02 wt%. the sample contains 2.7 wt% carbon. mf102,5_11cm176486x16oililmamhund_c.eps fe grms ilm 0.1 mm arm fe fe ru trofig. 110. reflected-light image of oxides and native iron in dacite, same sample as in fig. 109a. an ilmenite microphenocryst (ilm) is affected by progressive reduction and sulfidation and has partly recrystallised to rutile (ru), armalcolite (arm), troilite (tro) and native iron (fe). it is surrounded by the dacitic groundmass (grms), also with native iron. see detailed description in pedersen (1981). sample 176486, clast in conglomerate north of hammer dal, north-west disko. 128128 structure due to the mixing and mingling of more and less contaminated magma, as exemplified by sample 326550 (fig. 109d; table 6, pages 164–168). distribution. the dacites are restricted in occurrence to western and north-western disko from just north of mellemfjord to north of hammer dal and jamma at the coast (figs 108c, d). rhyolite composition and petrography. the rhyolitic rocks contain 70.5–76.9 wt% sio2; they include two samples that plot just inside the dacite field in a tas diagram (le maitre 2002) but are closely related to the rhyolites. the rhyolites are low in mgo (1.2–0.2 wt%) and tio2 (0.67–0.04 wt%). they comprise two compositional and petrographic groups both of which are peraluminous. these are garnet rhyolite, which is strongly porphyritic and sometimes transitional to dacite, and sanidine rhyolite, which is much less porphyritic and more evolved and has higher sio2 contents. because no ferromagnesian silicate or oxide mineral analyses have been published from the rhyolites, a selection of pyroxene, garnet, biotite, ilmenite and hercynite analyses is presented here in table 5. garnet rhyolite. these rocks, exemplified by sample 156518 (fig. 111a), consist of 65–70% glassy groundmass, c. 25% crystals and c. 5% xenoliths. the crystals comprise phenocrysts of quartz, plagioclase, orthopyroxene, biotite and almandine garnet, cognate clusters of plagioclase, orthopyroxene, pigeonite and garnet (fig. 111b), minor ilmenite, apatite, zircon and graphite, and various xenocrysts of sillimanite (fig. 111c), corundum, hercynitic spinel and feldspar. there is a variety of sediment xenoliths dominated by equilibrated mudstone (fig. 111a) and also some highly equilibrated clasts, perhaps of sandstone, with intergrown quartz, garnet and plagioclase (fig. 111d). graphite is as a significant component of many xenoliths and xenocrysts. at a locality described below, garnet ryholite clasts were picked up by a subaerial basalt lava flow that engulfed a conglomerate bed. these rhyolite clasts are rounded, weathered and oxidised; pyrometamorphism by the basalt magma has transformed the biotite, orthopyroxene, garnet and ilmenite into ferric-oxide-rich oxide-silicate assemblages (fig. 111e), while the original textures are preserved. in particular, the very characteristic quartz-garnet intergrowth patterns are now preserved as pseudomorphs consisting of quartz, iron oxide and silicate. sanidine rhyolite. this rock type mainly forms dark, glassy pitchstone (fig. 112a) but also microcrystalline laminated rocks. the glassy sanidine rhyolite contains c. 95% glassy groundmass, about 5% phenocrysts of quartz, plagioclase, sanidine and biotite (fig. 112b), and very minor ilmenite, zircon, apatite, monazite and almandinerich garnet. the rocks also contain very scarce xenocrysts of aluminous phases and highly equilibrated sediment xenoliths, which are now aggregates of feldspar, hercynite, biotite, orthopyroxene, ilmenite, sillimanite and graphite. the graphite occurs in xenoliths, enclosed in ilmenite and biotite, and as small flakes in the groundmass glass (fig. 112c). the microcrystalline sanidine rhyolites are strongly flow-laminated rocks with 80–85% groundmass and <20% phenocrysts of quartz, plagioclase, sanidine and biotite. the minor phases have not been investigated. the rocks are carbon-poor (organic c <0.02%) and graphite may have been oxidised away completely. facing page: fig. 111. thin sections (scanned) of garnet rhyolites of the nordfjord member. a: garnet rhyolite with phenocrysts of feldspar (fsp), quartz (qz), biotite (bi), orthopyroxene (opx) and garnet (not in image). the upper right half is a graphite-rich, magma-modified mudstone xenolith (xen with graphite). sample 156518, clast in conglomerate in sedimentkløften, hammer dal, north-west disko; sio2 = 72.6 wt%, tio2 = 0.49 wt%, mgo = 0.56 wt%. b: cognate cluster of garnet (gt), quartz (qz) and plagioclase (pl) in a glassy groundmass (grm). garnet rhyolite sample 113508, clast in conglomerate in sedimentkløften, hammer dal, north-west disko. c: sillimanite xenocryst (sil) and quartz phenocrysts (qz) in garnet rhyolite sample 113508, clast in conglomerate in sedimentkløften, hammer dal, north-west disko. d: the upper right half of this image is a cluster of highly equilibrated clasts, perhaps of sandstone, with intergrown quartz, garnet and plagioclase (gt+qz+pl) and interstitial smectite. the lower left half is the host rock, a garnet rhyolite with plagioclase (pl), quartz (qz), orthopyroxene (opx) and garnet (gt). sample 326468, clast in sedimentkløften, hammer dal, north-west disko. e: a former garnet rhyolite pebble that has been picked up in a basalt lava flow and metamorphosed at high temperature and low pressure. garnet-quartz-plagioclase aggregates have recrystallised to iron oxide-quartz-plagioclase aggregates (feox+qz+pl), whereas the rhyolite textures with groundmass (grm), plagioclase (pl) and quartz (qz) are preserved. sample 274424, hammer dal, north-west disko; sio2 = 72.6 wt%, tio2 = 0.47 wt%, mgo = 0.84 wt%. 129 mf103a_8cm156518xenutuschbmodcut_c.eps a pl fsp opx bi qz xen with graphite 5 mm mf103b_8cm113508x4granatogqzmodcut_c.eps b pl grm gtqz 5 mm mf103c_8cm113508x4sillimanitcut_c.eps c silsil qz 1 mm mf103d_8cm326468xpolmodcut_c.eps d gt+qz+pl gt+qz+pl gt qz pl opx 10 mm mf103e_8cmtværs274424pol5modcut_c.eps e grm pl qzfeox+qz+plfeox+qz+pl 10 mm 130130 xenoliths in rhyolite. mudstone xenoliths in the nativeiron-bearing rocks elsewhere on nuussuaq and disko comprise a suite of more or less equilibrated rocks, from melted but unequilibrated buchites composed of glass, cordierite, mullite, sulfide and graphite to highly magmaequilibrated rocks with plagioclase, spinel, corundum and graphite (e.g. törnebohm1878; pedersen 1979a; pedersen & larsen 2006). the equilibrated rocks have exchanged elements with the basic magma, and a marked depletion in some elements and a strong enrichment in others, e.g. ca, has taken place. of the xenoliths in rhyolite, a single c. 3 cm large, equilibrated mudstone xenolith has been analysed (see table 13, sample 156518.2). this shows that the reaction between the mudstone and the rhyolite magma was different from the reaction with the basic magma: the mudstone has taken up k and ba from the rhyolite to stabilise biotite and k-feldspar, whereas the uptake of ca is small. the xenolith now contains 6.15 wt% k2o, which is much more than measured in any sediment in the nuussuaq basin, and has retained much of its carbon as graphite (11.72 wt% c). distribution. rhyolite as glassy and microcrystalline massive rock has only been found as transported cobbles and pebbles in conglomerates on west disko from vesterdalen north-east of hammer dal, and in a river deposit on north-east disko (described below). however, rhyolite is extensively distributed as primary airfall tuffs and altered and redeposited tuffaceous sediments. 1 2 3 4 5 6 7 8 9 10 11 12 156518.1 156518.1 113508 156518.1 113508 113515 156518.1 113515 156518.1 113515 156518.2 156518.2 pigeonite orthopyx orthopyx garnet garnet garnet biotite biotite ilmenite ilmenite hercynite hercynite 49.43 46.61 47.71 37.51 36.86 37.21 33.28 33.47 0.39 0.22 0.13 0.14 0.24 0.12 5.72 4.90 50.01 51.76 0.44 1.85 1.09 1.42 0.98 22.28 21.48 22.15 14.78 16.44 0.23 0.29 59.33 37.53 0.16 12.80 30.12 43.89 44.35 32.16 35.19 36.90 29.14 28.97 46.86 46.17 34.15 42.11 0.55 0.46 0.73 0.62 1.23 2.16 0.35 0.53 0.25 0.23 13.06 6.24 6.60 6.00 2.99 2.20 4.63 4.27 0.80 0.63 5.67 2.14 5.11 0.77 0.52 1.51 2.62 1.65 0.11 0.17 0.24 0.54 0.58 0.74 1.03 8.81 8.90 99.99 100.15 101.60 100.22 100.61 102.39 97.21 97.98 98.25 99.38 100.00 96.83 1.937 1.928 1.944 2.946 2.946 2.943 5.375 5.330 0.000 0.000 0.000 0.000 0.050 0.069 0.047 2.063 2.024 2.065 2.814 3.086 0.007 0.009 1.950 1.404 0.011 0.007 0.004 0.008 0.014 0.007 0.695 0.587 0.957 0.982 0.009 0.044 0.004 0.321 0.987 1.518 1.511 2.112 2.352 2.441 3.936 3.858 0.998 0.974 0.796 1.118 0.018 0.016 0.025 0.041 0.083 0.145 0.000 0.000 0.008 0.011 0.006 0.006 0.763 0.385 0.401 0.702 0.356 0.259 1.114 1.013 0.030 0.024 0.236 0.101 0.215 0.034 0.023 0.127 0.224 0.140 0.019 0.006 0.018 0.043 0.046 0.232 0.318 1.815 1.808 43.59 20.21 20.96 24.95 13.15 9.60 22.07 20.80 2.95 2.37 22.83 8.30 4 4 4 8 8 8 16 16 2 2 3 3 no. ggu no. mineral oxides, wt% sio2 tio2 al2o3 cr2o3 feo* mno mgo cao na2o k2o sum recalculated to cations no of cations si al ti cr fe2+ mn mg ca na k mg number 1. pigeonite in glomerocryst in garnet rhyolite 156518.1 2. orthopyroxene margin on cognate pyroxene-feldspar cluster in garnet rhyolite 156518.1 3. orthopyroxene in pyroxene-garnet aggregate in garnet rhyolite 113508 4. garnet fragment in glass in garnet rhyolite 156518.1 5. garnet in garnet-ilmenite aggregate in garnet rhyolite 113508 6. small garnet clast in glass in sanidine rhyolite 113515 7. biotite in glass in garnet rhyolite 156518.1 8. biotite in glass in sanidine rhyolite 113515 9. ilmenite intergrown with biotite in garnet rhyolite 156518.1 10. ilmenite in glass in sanidine rhyolite 113515 11. hercynite in magma-modified mudstone xenolith 156518.2 in garnet rhyolite 12. chromian hercynite in magma-modified mudstone xenolith 156518.2 in garnet rhyolite analyses of bulk rocks 156518.1 and 156518.2 are shown in tables 6c and 13, respectively. feo* is total iron as feo. mg number is 100 × mg/(mg+fe). 156518.2 is a xenolith in sample 156518.1 when recast to equivalent oxygen, the pyroxenes, garnets and oxides show no significant fe2o3, with one exception: the chromian hercynite no. 12 with 0.2 fe3 cations. the oxidation state of biotite cannot be calculated because of possible vacancies in the crystal lattice. table 5. microprobe analyses of minerals in rhyolites of the nordfjord member 131 geological themes and locality descriptions sediment horizons at the base of the nordfjord member the base of the nordfjord member is marked by a prominent sediment horizon deposited on the eroded top surface of basaltic lavas of the upper rinks dal member. depending on local topography the sediment may exceed 2 m in thickness but may also be just a few decimetres thick (fig. 113). on north-western disko, the lower part of the sediment horizon is a yellowish-brown claystone which sometimes contains beds of extensively weathered volcanic material interpreted as derived from rhyolitic pumice tuffs, sometimes with graphite. this indicates that formation of strongly sediment-contaminated and evolved rhyolitic magmas took place from the very beginning of the formation of the nordfjord member, in this case even before the beginning of the basaltic to andesitic activity (several similar events took place later). three particularly significant localities are described below. niaqussat (fig. 103, loc. 1; fig. 14, profile 12; fig. 113, profiles c, d). the locality is shown from a distance in fig. 114. the eroded basalt lava surface of the upper rinks dal member (unit 517) is overlain by a few metres of claystone and tuffaceous sandstone with strongly altered rhyolitic pumice. this sediment is covered, partly by a subaerial basalt lava flow, and partly by c. 10 m of stratified conglomerate deposited by a river, which eroded laterally into the lava flow (fig. 113). the lava flow and the conglomerates are in turn covered by two basaltic lava flows with prominent colonnades and entablatures indicative of emplacement over a wet surface (fig. 115). the conglomerate comprises three beds. the lowest bed has a matrix of claystone and mostly carries clasts of basalt from the rinks dal member but also a minor component of highly weathered clasts of picrite, indicating that parts of the vaigat formation must have been exposed to erosion at that time. the upper two conglomerate beds likewise carry basalt clasts but also clasts of rhyolitic pumice and pitchstone (fig. 115), and parts of the matrix are rich in quartz crystals and decomposed rhyolite pumice (fig. 116). altogether, the basal conglomerates of the nordfjord member at niaqussat indicate tectonic activity between the deposition of the rinks dal and nordfjord members. mf104a_8cm156516polukumodl_c.eps a 5 mm glass bio qz fsp mf104b_8cm156559x4biqzplirh_c.eps b 5 mm bio bio glass qz san mf104c_8cmtværs113515x4biotilmgrafirhyreflmod3_c.eps c 0.4 mm glass bio zr ilm fig. 112. thin sections of sanidine rhyolites of the nordfjord member. a: glassy pitchstone with about 5% phenocrysts of quartz (qz), feldspar (fsp) and biotite (bio). sample 156516, clast in conglomerate in sedimentkløften, hammer dal, north-west disko; sio2 = 76.6 wt%, tio2 = 0.10 wt%, mgo = 0.17 wt%. plane polarised light. b: cluster of phenocrysts of quartz (qz), sanidine (san) and biotite (bio) in glass in sanidine rhyolite sample 156559, clast in conglomerate north of rink dal, north-west disko. plane polarised light. c: reflected-light image of sanidine rhyolite showing flakes of graphite (at yellow arrows) enclosed in ilmenite (ilm), in biotite (bio) that forms a reaction rim on ilmente, and in the groundmass glass. note also minor zircon (zr) in biotite. sample 113515, clast in conglomerate in sedimentkløften, hammer dal, north-west disko. 132132 the type section of the nordfjord member north of hammer dal (fig. 103, loc. 5; fig. 104; fig. 14, profile 11; fig. 113, profile b). this section includes two sediment horizons with a lava flow between them near the lower boundary (fig. 104 around 1800 m). the lower sediment horizon is a 10–20 cm thick, yellowish-brown claystone; the chemical composition of the overlying lava flow suggests, without being definitely diagnostic, that it belongs to the upper rinks dal member (unit 518). the base of the nordfjord member is therefore placed at the base of the upper sediment horizon. this is a 1.6 m thick volcaniclastic sediment composed of several units of which the lower c. 45 cm is a lateritic soil with scattered volcanic clasts (fig. 113). the soil is covered by a c. 40 cm thick, matrix-rich conglomerate bed with up to 4 cm large clasts of basalt and, remarkably, clasts of garnet rhyolite and scarce millimetre-sized clasts of almandine garnet. the matrix is partly derived from a rhyolitic source. the conglomerate is covered by c. 60 cm of claystone, greenishgrey and reddening upwards, with altered rhyolite clasts containing recognisable phenocrysts of quartz and biotite. the 10–15 cm thick top zone consists of tuff and lateritic soil that has been strongly baked by the overlying lava flow. again, it is evident that highly evolved rhyolitic rocks were available as a source for the initial sediments at the base of the nordfjord member. 3 m 2 0 1 278 280 282 284 286 288 290 292 278 280 282 284 286 288 290 292 2 m 0 1 no exposure photo: fig. 116 no exposure alkali basalt baked baked urdm urdm urdm a b c d niaqussat east west nordfjord point 1070 m hammer dal point 440 m northern gully gv04_06_034_lml.eps fig. 113. detailed logs through the sediment horizon at the base of the nordfjord member. note different vertical scales. for legend, see fig. 117, for localities, see fig. 103, locs 1, 5, 22. 133 the point 1070 m section north of nordfjord (fig. 103, loc. 22; fig. 14, profile 8; fig. 113, profile a). this section shows the well-exposed base of the nordfjord member. the strongly weathered surface of a lava flow of the upper rinks dal member (unit 518) is overlain by a 2.5 to c. 4 m thick succession of claystone. the lower c. 2 m of the claystone is yellowish-brown and contains a number of strongly altered tuff beds and beds with altered volmf106_17cma9021_5_01mod2_c.eps 520 la 520 basaltic andesite 520 andesite 326659 326643 520 sedurdm urdm fig. 114. overview photo of the nordfjord member at niaqussat, north-west disko. the nordfjord member begins with a basal sediment (520 sed) that is overlain by lava flows (520 la) of basalt, basaltic andesite and a single native-iron-bearing andesite erupted from a nearby crater. some samples from profile 12 in fig. 14 are indicated. the circle indicates the location of log d in fig. 113 and figs 115, 116. urdm: upper rinks dal member. mf107_11cm1984_06_20mod_c.eps fig. 115. the basal c. 12 m thick sediment horizon of the nordfjord member at niaqussat (fig. 113, log d). the locality is encircled in fig. 114. the sediments comprise claystone, tuffaceous sandstone and bedded conglomerates described in the text. the prominent colonnade of the overlying lava flow indicates a wet environment during its emplacement. 134134 canic clasts and remnants of plant fossils. the upper half of the claystone gradually becomes more reddish, and the uppermost 20 cm is a red lateritic soil, which is strongly baked by the overlying alkali basaltic lava flow. the west disko graphite rhyolite (wdgr) volcano the west disko graphite rhyolite volcano (wdgr) denotes one or several volcanoes located somewhere in the westernmost part of north-western disko or on the shallow shelf offshore disko. the volcano erupted rhyolitic tuffs but also massive glassy lavas and perhaps bombs. one of the unusual features of the erupted rocks is the ubiquitous presence of graphite, which, together with scattered magma-modified sediment xenoliths, shows that the rhyolitic magmas evolved in high-level reservoirs within carbonaceous mudstones and sandstones. the main occurrences of wdgr rocks are situated on north-western disko. they were first noted by ggu geologists in a small gully in the south wall of hammer dal (v. münther, personal communication 1968); the gully was subsequently named sedimentkløften (‘sediment gully’). they were described by pedersen (1969, 1975a, 1977a), pedersen & pedersen (1987) and hansen & pedersen (1985). mf108_8cmakp_1984_06_16mod_c.epsfig. 116. clast-supported conglomerate in the basal sediment horizon of the nordfjord member at niaqussat. the rounded clasts are mainly basalt but clasts of rhyolitic pumice also occur. the matrix is rich in quartz and decomposed rhyolite pumice. for location of the photo, see fig. 113, log d. length of ruler 15 cm. lava flow, basalt to basaltic andesite lava top, entablature and top rubble legend (facing page, figs 113, 124) strong thermal metamorphic overprint white to yellow, rhyolitic pyroclastic beds: clasts in pyroclastic beds: grey, rich in disseminated graphite, rhyolitic altered to clay intermediate to basaltic reworked cross-bedded conglomerate; clasts of basalt (blue), picrite (green), rhyolite (yellow), garnet rhyolite (yellow + red dot), graphite rhyolite (grey), claystone (brown), sandstone (white), laterite (red) euhedral quartz phenocrysts volcaniclastic sandstone and conglomerate (river deposit) lithic clasts, non-rhyolitic lithic clasts, rhyolitic (pitchstone, garnet rhyolite) pumice clasts of white rhyolite pumice clasts of grey graphite rhyolite coalified plant debris rhyolitic glassy rock formed by remelting of tuff or volcaniclastic rock lateritic soil, purple to red claystone, reddish brown claystone, dark brown claystone, yellowish brown gv04_06_032_lml_legende.eps facing page: fig. 117. detailed logs of three successions with rhyolite tuffs on north-west disko. locality numbers are those in fig. 103. unreworked, primary airfall tuffs are only found north of hammer dal. note the different vertical scales in the sedimentkløften and north of hammer dal logs. several photographs are located on the logs. the legend also covers figs 113 and 124. 135 2.5 m 2 1.5 1 0.5 0 m 4.5 4.9 4 3.5 3 altered basalt scoria 7 tuff layers erosion graphite-rich + garnet matrix rich in rhyolite pumice basaltic matrix ph ot o, fi g. 12 0 photo, fig. 123a photo, fig. 122b photo, fig. 122a photo, fig. 123b photo, fig. 119 sandstone xenoliths basaltic andesite gv04_06_033_lml.eps 2.5 m 2 1.5 1 0.5 0 no exposure 5 loc. 2 loc. 12 loc. 5 north of hammer dal sedimentkløften 4.5 4 3.5 3 2.50 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 m 7 m 6.5 6 5.5 136136 the area on north-western disko with the main occurrences of wdgr rocks is generally poorly exposed and extensively faulted, and large areas along the coasts and in the valleys are covered by quaternary deposits. because of the unique nature of the rocks, and in order to locate the volcano or lava flows from the volcano, all the gullies and valley sides in the western part of the area between giesecke dal and nordfjord have been systematically searched. while many individual exposures of wdgr tuffs and coarser volcaniclastic beds were located, no feeders, craters or rhyolitic lava flows from the wdgr volcano have ever been found. timing and duration of the wdgr volcanism. tuffs and conglomerates with wdgr rocks occur in the volcaniclastic deposits at the base of the nordfjord member, immediately overlying the rinks dal member, and the volcano must therefore have developed in the time interval represented by the basal sediment horizon. on the other hand, there are no traces of primary wdgr rocks within the uppermost part of the nordfjord member, and the volcano thus appears to have become extinct before the formation of this member ended. tuff successions tuffs derived from the wdgr volcano occur on western and north-eastern disko, but undisturbed tuffs many metres thick are only preserved in a limited area on north-western disko from rink dal to north of hammer dal (fig. 103). redeposited tuffs and conglomerates in sedimentkløften. an about 18 m thick succession of volcaniclastic sandstone, conglomerate and lava flows is exposed within sedimentkløften in the south wall of hammer dal. the strata dip 30°w and are disturbed by faulting. the succession belongs to the lower part of the nordfjord member but is not stratigraphically well constrained. it contains the most diverse assemblage and the largest rhyolite blocks found and is probably the exposure that is most proximal to the volcano. it is reconstructed here as a vertical log in fig. 117. at the base of the succession, a weathered basalt flow with a surface of basalt blocks in a claystone matrix is covered by 2 m of claystone and rhyolite tuff. the most conspicuous unit is the following, c. 5 m thick, light grey-white tuffaceous sandstone which is visible from a distance of several kilometres and has given the gully its name (fig. 118). this sandstone contains rounded pebbles and cobbles of garnet rhyolite and must have been derived by reworking of material produced by explosive activity associated with one of the main rock types of the wdgr volcano, the garnet rhyolite (fig. 119). it is covered by a sediment-contaminated basalt lava flow with sandstone xenoliths. on top of the flow is a c. 5 m thick bedded succession composed of reworked rhyolite tuff, a mf110_17cmakp_1972_h_01_06sedklhdmod_c.eps fig. 118. the tuffaceous sediment succession in sedimentkløften; a log is shown in fig. 117. note encircled person for scale. 137 conglomerate with a large diversity of clasts up to boulder size ranging from several types of rhyolite to basalt and alkali basalt, and an upper oxidised conglomerate with clasts of basalt, laterite and basaltic scoria (fig. 120). airfall tuffs north of hammer dal. the three best preserved tuff localities occur in the hammer dal area, two north of hammer dal and a third south of hammer dal (fig. 103, locs 2, 5, and 27). the tuffs have minimum thicknesses of 7.1 m, 4.85 m and 5.3 m, respectively. detailed logs of the successions at the two northern localities are shown in fig. 117. figure 121 is an overview photo of tuff locality 5. the most distinctive constituent of the tuffs, which is recognised in all profiles through airfall tuffs, is a c. 20 cm thick horizon of bedded, graphite-rich rhyolitic tuffs (fig. 122). graphite gives the pumice a dark grey to black colour. the horizon comprises five to six thin, fine-clastic beds, which vary in colour between light grey and black, followed by seven coarser pumice beds, some of which are mixtures of white rhyolite pumice and dark grey graphitic rhyolite pumice. the horizon is overlain by a white-grey rhyolitic pumice bed (fig. 122a), which is more than 1 m thick at the two northern localities but only 30 cm at the southern locality. hand-picked separates of a white and a dark grey pumice clast from a mixed pumice layer were analysed for toc and sulfur (see table 7, sample 326497). the white clast contains 0.51 wt% toc whereas the dark grey clast is graphite-rich and contains 3.1 wt% toc; mf111_8cmakp_1968_sed1hdmdglmod_c.eps 1 m fig. 119. rhyolitic tuffaceous sandstone in sedimentkløften. the two red arrows point to garnet rhyolite clasts. for location of the photo, see the log in fig. 117. mf112_17cmakp_1972_h_01_07sedklhdmod2_c.eps 1 m tu�-sed cgl ox-cgl basalt fig. 120. the upper part of the tuffaceous sediment succession in sedimentkløften: a bedded succession of reworked rhyolite tuff (tuff-sed), a light-coloured conglomerate (cgl) and an upper oxidised conglomerate (ox-cgl), overlain by a basalt lava flow. the succession dips towards the viewer. the boulder at the hammer has rolled down from the light conglomerate; it is the largest rhyolite boulder found on disko. length of hammer 47 cm. for location of the photo, see the log in fig. 117. 138138 the difference demonstrates a significant heterogeneity in the erupting high-level magma chamber. in the thick tuff succession at locality 2, the lower 3.8 m are dominated by a number of claystone beds varying in colour between greyish-green and dark grey, which contain altered clasts of rhyolitic pumice. notably, a number of these beds contain clear bipyramidal quartz crystals (phenocrysts) that are perfectly preserved, whereas their host pumice has decomposed to clay. together, the thick tuffs at the two northern localities have preserved evidence of 24 explosive eruptive events, with many more small, individual beds; 20 events were rhyolitic and three non-rhyolitic. graphite is important in four of these eruptive events. plant fossils occur at four to five levels, and clay and lateritic soil at six to eleven levels. loose trunks of silicified wood occur at locality 2 and must have been eroded out of the tuff succession. these occurrences indicate significant time gaps between the individual eruptions of the tuffs. moreover, there are signs of local erosion and reworking by water of several tuffs (fig. 123), notably in the southern locality 27, which comprises two prominent conglomerate beds. the intensive weathering to clay of the lower part of the succession and the better-preserved state of the upper 3 m of the tuffs indicate a considerable duration of the explosive eruptive phase and a much higher eruption rate for the uppermost 3 m of the tuffs. altogether, the distribution of the airfall tuffs indicates a location of the wdgr volcano around or just north of hammer dal or within the neighbouring offshore area. conglomerates and volcaniclastic sandstones in addition to the conglomerates and sandstones described above, the nordfjord member contains a number of other conglomerate and volcaniclastic sandstone beds which all indicate active tectonic movements, recurrent phases of local erosion, and the presence of rivers. the main occurrences of such sediments are on western and north-western disko from nordfjord to giesecke dal and in a limited area on north-eastern disko. they seem to be absent on south-western disko, and no evidence of conglomerate beds is found in the large areas where only erosional remnants of the nordfjord member are preserved. in the poorly exposed and severely block-faulted western and north-western disko, individual sediment beds are only preserved over short lateral distances from a few metres to at most a few hundred metres, and no reconstructions of palaeo-riverbeds are possible. in contrast, a mf113_11cm1984_03_27l84-19modcut_c.eps 5 m fig. 121. tuff horizon north of hammer dal, loc. 5 in fig. 103. a detailed log is shown in fig. 117. 139 0.5 m mf114a_8cma1974_lok7432_surtufno28mod_c.eps a mf114b_8cm1984_04_22l84-27mod_c.eps b fig. 122. a characteristic horizon of bedded, graphite-rich rhyolite tuff found in all profiles through airfall tuffs, here at locality 2 north of hammer dal. for description, see text. a: overview of the succession with the c. 15 cm thick, graphite-rich tuff horizon in the centre. b: close-up of the graphite-rich tuff horizon. the thin tuff layers (1–2 cm) in the lower part show normal grading, which suggests that they may have been deposited in water. the absence of wave or current-generated cross-lamination suggests rapid deposition in a low-energy environment. the thick, coarse-grained tuff layer in the middle indicates either a more violent eruption or a crater much closer to the locality. for location of the photos, see the log in fig. 117 (where the measured tuff layer is 20 cm thick). mf115a_8cmakp_1984_06_30l84-47modcut_c.eps a mf115b_8cmakp_1984_03_36l84-19mod_c .eps b fig. 123. rhyolite tuffs north of hammer dal showing signs of reworking by water. a: well-sorted rhyolite tuff with cross-bedding indicating fluvial transport and deposition. b: moderately sorted, very coarse-grained tuffaceous sandstone with weak stratification, interbedded with two thin dark claystone layers in the upper part of the photo, suggesting settling from suspension in water. for locations of the two photos, see the logs in fig. 117. 140140 0 1 2 3 4 5 0 1 2 3 4 5 m 6 7 8 m 9 10 11 12 13 14 15 16 no exposure disturbed by permafrost base of invasive lava flow top of invasive lava flow m below invasive lava flow above invasive lava flow gv04_06_036_lml fig. 124. log of c. 21 m of rhyolitic tuffaceous sediments in a river deposit, north-eastern disko. the succession is split in two by an invasive lava flow. legend in fig. 117. three sedimentary facies are distinguished: mudstone, presumably altered tuff; pebbly sandstones with weak stratification, locally interbedded with thin mudstones; and conglomerates, massive or weakly cross-bedded. the mudstones are interpreted as floodplain deposits and the sandstones and conglomerates are fluvial. see text for descriptions and discussion. 141 horizon of conglomerate and sandstone on north-eastern disko is continuous over at least 3 km. the conglomerate and sandstone beds contain very different clast populations. here two types will be described, namely beds with a substantial component of rhyolitic rocks from the wdgr volcano and beds without rhyolitic rocks. sediments with clasts from the wdgr volcano river deposit on north-eastern disko. the most extensive conglomerate-bearing horizon within the nordfjord member is an up to 20 m thick deposit of tuffaceous sandstone and conglomerate rich in rhyolitic clasts and pumice, which is exposed in the steep coastal cliff below point 1530 m near uunartuarsuk on north-eastern disko (central disko section at 69 km; fig. 10, profile 6 at 1315−1365 m; sediment log in fig. 124). the deposit forms a white sediment layer that can be followed for 3 km along the cliff face. no similar sediments are present in the surrounding areas, and the massive rhyolite clasts cannot have been deposited by airfall from a volcanic explosion in the wdgr volcano more than 75 km to the north-west. this, and the rounded and abraded state of the clasts, indicate that the sediment was deposited by a river, which must have transported rhyolitic rocks from the wdgr volcano on north-western disko for at least 75 km towards south-south-east across northern disko (larsen & pedersen 1989). the deposit is only accessible at the eastern end of the exposure where it is partly disturbed by a c. 30 m thick invasive lava flow of the nordfjord member (fig. 125; also larsen & pedersen 1989, figs 2 and 4). the flow has intruded the upper part of the succession and has melted the sediments at its lower and upper contacts into glassy zones in which reddish layers of oxidised claystone can mf117_17cmakp_1985_09_08l85-30mod_c.epsfig. 125. few metres of rhyolitic, tuffaceous sediment deposited on the irregular, eroded top of the underlying lava flow. the sediment is relatively fine grained with parallel bedding and is baked (dark grey colour) by the overlying, invasive lava flow. the lower boundary of the tuff layer is mostly covered by down-washed tuffaceous material. nordfjord member below point 1530 m, north-eastern disko. the part of the invasive flow seen here is c. 20 m thick. see also fig. 10, profile 6. 142142 still be distinguished from blue-grey layers of melted volcaniclastic sediment rich in rhyolitic pumice (fig. 126). close to a palaeo-riverbank cut into the lowermost lava flow of the nordfjord member, there are c. 15 m of sediment below the invasive lava flow and more than 5 m above the flow (fig. 124). the lowermost c. 3.5 m of the sediment consist of yellowish-brown to dark-brown claystone beds with a few millimetre-thick layers of coalified plant remains. this is interpreted as a floodplain deposit. upwards, the succession is composed of alternating beds of volcaniclastic sandstone and conglomerate, in places separated by thin layers of yellowish brown claystone. numerous erosive structures are present (fig. 126). the volcaniclastic beds have a whitish-grey matrix dominated by reworked rhyolitic pumice (fig. 126). the conglomerate comprises basalt clasts up to 15 cm large and rhyolite and andesite clasts up to 5 cm large. among the rhyolites, which all carry graphite, are the three characteristic components of the wdgr volcano: glassy quartz-biotitefeldspar rhyolite (pitchstone), which is predominant; flow-laminated microcrystalline rhyolite (felsite, from lavas) with the same phenocrysts, which is common; and scarce, strongly porphyritic garnet rhyolite. andesites, and possibly dacites, with xenoliths of magma-modified mudstone are also present. the main part of the sediment is transported from north-western disko, and the many beds separated by claystone show that the succession was deposited over some time and not in a single catastrophic event. conglomerate deposits on north-western disko. it is not possible to correlate any conglomerate bed on western disko with the fluvial deposit on north-eastern disko, although some similarities are seen in a conglomerate mf118_11cmakp_1985_09_17l85-30mod_c.eps fig. 126. sediments in rhyolitic tuffaceous deposit on north-eastern disko. pebbly sandstone and fine-grained sandstone form trough-shaped sets. the parts with contrasting grain sizes are well separated, indicating abrupt changes from high to lower energy conditions; deposition could have taken place in a fluvial channel. the base of an invasive lava flow is seen in the upper part of the photo; the sediment at the contact has been melted into a glassy rock; the glassy interval is indicated by a white vertical line at upper right; the amount of glass decreases downwards from the contact. in the glassy interval, a reddish layer of oxidised claystone (a former soil horizon?) can be distinguished from blue-grey layers of melted volcaniclastic sediment rich in rhyolitic pumice. length of field spade 67 cm. 143 horizon about 70 m above the base of the nordfjord member in the type section (fig. 103, loc. 5; fig. 14, profile 11, lower part at 173 m), and in a parallel gully. this horizon is 10.2 m thick and composed of volcaniclastic sandstone and conglomerate rich in rhyolite clasts, with a matrix that is light grey due to a large component of reworked rhyolitic pumice. the clast size varies from pebbles to rare boulders. the clasts vary considerably in composition from bed to bed and from locality to locality, but strongly sediment-contaminated volcanic rocks are always abundant and particularly rhyolitic rocks from the wdgr volcano. a typical fresh conglomerate surface is shown in fig. 127. the rhyolitic rocks comprise pitchstones, felsites and garnet rhyolites just as in the north-east disko conglomerates. dacites with native iron are common, as are basaltic to andesitic rocks. a characteristic but rather scarce rock type, which is only known from these conglomerates is graphite-rich high-al dacite. at some localities feldspar-phyric basalt clasts seem to be missing, at others they are subordinate. pitchstone glass clasts are very common to dominant in the sand to gravel fraction. south of hammer dal there is a conglomerate bed with a clast population dominated by basalt, but also mf119_11cma1984_04_12l84-20mod_c.eps pu pt pt bas fig. 127. conglomerate bed north of hammer dal, north-west disko. matrix-supported, poorly sorted conglomerate bed with rounded to sub-angular clasts. the clasts comprise basalt (bas), pitchstone (pt) and white rhyolitic pumice (pu). the light grey matrix has a large component of reworked rhyolitic pumice. the conglomerate may be a debris-flow deposit. for locality, see fig. 103, loc. 5. fig. 128. conglomerate bed south of hammer dal, north-west disko. matrixsupported, poorly sorted conglomerate bed, tentatively interpreted as a debris flow deposit. the clasts comprise basalt, garnet rhyolite, felsite, pitchstone and rare clasts of picrite derived from the vaigat formation. length of hammer 32 cm. for locality, see fig. 103, loc. 13. mf120_11cmakp_1985_09_31l85-32mod_c.eps 144144 with common clasts of garnet rhyolite, felsite and pitchstone (fig. 128). a rare component is weathered and well-rounded clasts of picrite, which must have been derived from the vaigat formation. conglomerate engulfed by a lava flow. about 2.8 km northeast of point 440 m north of hammer dal (fig. 103, loc. 3), a more than 6 m thick feldspar-phyric basalt lava flow of the nordfjord member has engulfed a conglomerate with rhyolitic blocks from the wdgr volcano. the lower 2 m of the flow are very rich in rounded xenoliths of several types of basalt and rhyolite and show chill zones against the blocks (fig. 129). some of the blocks were weathered before they were picked up by the flow. during reheating in the flow, pumice clasts were partly remelted and the almandine garnets of the garnet rhyolites were transformed into oxide-silicate assemblages, which have preserved the original characteristic garnet-quartz intergrowth texture of this rock type (fig. 111e). sediments without clasts from the wdgr volcano a number of uncorrelatable exposures of metre-thick conglomerates and volcaniclastic sandstones without rhyolitic clasts occur in an area on north-west disko from just south of the western part of hammer dal, through the western part of rink dal to the north coast of nordfjord. the sediments are sandwiched between plagioclase-phyric basalts within the upper part of the nordfjord member. the most extensive of these sediment deposits is found in the northern and southern walls of the western part of rink dal (fig. 130; fig. 103, loc. 15). the succession is up to c. 10 m thick and composed of beds of volcaniclastic sandstone alternating with variably coarse conglomerate beds. the clasts vary in size from less than 1 cm to more than 2 m and predominantly consist of plagioclasephyric and aphyric basalt; in addition, there are strongly sediment-contaminated basaltic andesites and andesites, some of which contain native iron and graphite-rich modified mudstone xenoliths. native-iron-bearing lava flows are not preserved within this part of rink dal, but they are present to the north, south and east. within a few kilometres from the main conglomerate locality in the north wall of rink dal, there are several localities with cross-bedded volcaniclastic sandstones, one of which contains a rich assemblage of fossil leaf imprints and coal fragments (fig. 103, loc. 14). the sandstones are covered by a decimetre-thick layer of lateritic soil, and at one locality there are two volcaniclastic sandstone beds separated by lateritic soil and an additional lateritic soil layer on top (fig. 131). this demonstrates that the small stream that deposited the sand dried up several times before the sand was covered by lateritic soil and finally by the next basalt lava flow. simple dacite lava flows between hammer dal and jamma one or two dacite lava flows with native iron occur in the strongly faulted and poorly exposed area that extends from hammer dal and about 10 km northwards to jamma (fig. 108d). the mineralogy and textures are interpreted as recording a trend of progressive reduction in t-fo2 space, as described in detail by pedersen (1981). most localities have only preserved part of a single dacite lava, but the north wall of hammer dal at point 600 m shows a continuous exposure of the uppermost 240 m of the nordfjord member (fig. 14, profile 9; figs 104c, 132; also pedersen 1977a, fig. 7). here a c. 120 m thick dacite lava flow with native iron is exposed, the thickest recorded lava flow in the paleocene of the nuussuaq barhy mf121_8cma1984_07_03l84-16mod_c.eps basalt lava rhy fig. 129. conglomerate with rhyolitic clasts (rhy) engulfed by a nordfjord member basalt lava flow. north of hammer dal, north-west disko. for locality, see fig. 103, loc. 3. 145 mf122_17cma1972_01_11mod_c.epsfig. 130. succession of conglomerate and sandstone beds within the upper part of the nordfjord member. the conglomerates are moderately sorted and are overlain by red volcaniclastic sandstone. the sediments do not contain any rhyolitic component, and due to their predominantly basaltic composition, they are not visibly baked by the overlying lava flow. length of hammer 47 cm. rink dal, north-west disko. for locality, see fig. 103, loc. 15. mf123_11cma1984_05_37l84-38modcut_c.eps 520 la lat lat sst sst fig. 131. sediments within the nordfjord member. two volcaniclastic sandstone beds (sst) each overlain by a lateritic soil (lat), indicating alternating wet and dry conditions. the sediments were eventually covered by a lava flow (520 la). length of ruler 15 cm. north of rink dal, north-west disko. for locality, see fig. 103, loc. 14. 146146 sin. its lower part is very compact and has a rusty brown weathering colour; it is jointed into more than 1 m thick columns and blocks. the upper c. 20 m is a blocky lava top. part of the same flow is exposed in an e–w-running gully c. 1.5 km to the north (fig. 104b; fig. 14, profile 10) and there are other small exposures in gullies 3 km to the south in the south wall of hammer dal, but no realistic map of the extent of the lava flow can be made. no feeders or intrusive rocks of similar composition are known. a very similar rock (pedersen 1981, table 1 sample 176466) is preserved as wave-eroded skerries around the small point jamma at the coast (fig. 103). the skerries consist of a dark grey, very fine-grained dacite with native iron (table 6; pedersen 1981, table 1 and 2, sample 176471). it is not seen in contact with other volcanic rocks. in addition to the exposed native-iron-bearing dacite lavas in the upper part of the nordfjord member, nativeiron-bearing dacite also occurs as clasts in conglomerate beds with rhyolite (see above) embedded within basaltic lavas older than the dacite lavas (fig. 14, profile 11, lower part at 170 m). there are both native-iron-bearing dacite clasts that are very similar to the lavas (pedersen 1981 table 1 and 2, sample 176486) and clasts of graphite-rich, high-al dacite of a type not known as lavas and likely to be associated with the wdgr volcano (see above). composite lava flows basaltic andesites, andesites and dacites commonly form part of composite lava flows or groups of flows from the same eruption centre. as shown in the last chapter, they are likely to have erupted from composite feeder dykes. the distribution and extent of some of the more characteristic composite flows is shown in fig. 108. eqaluit/nordre laksebugt. a c. 27 m thick, composite lava flow at the base of the nordfjord member occurs 2.5 km east-north-east of the river outlet at eqaluit (fig. 133). the flow is only known from this locality. it has a 2.5–3.5 m thick lower part of columnar-jointed, nearly aphyric basaltic andesite with 53 wt% sio2, 2.2 wt% tio2 and 5.8 wt% mgo (table 6, sample 176579). this passes upwards through a thin hybrid zone with incomplete mixing into a more than 20 m thick, more silicic basaltic andesite with 55 wt% sio2, 3.1 wt% tio2 and 3.6 wt% mgo (table 6, sample 176582). this part is dotted with rusty spots of weathered native iron and troilite; it contains numerous xenoliths of graphitic, magma-modmf124_17cma1972_k600hamdalmod2_c.eps fe-graphite dacite fe-dacite 100 m fig. 132. a nordfjord member lava succession with two native-iron-bearing dacite flows, a lower 120 m thick and an upper 45 m thick, with a series of thinner andesite flows in between. point 600 m, north side of hammer dal, west disko, looking north. see also fig. 14, profile 9 and the photogrammetric interpretation in fig. 104c. for locality, see fig. 103, loc. 8. 147 ified mudstone. the upper 5 m of the flow is vesiculated and strongly oxidised. the flow provides an example of contamination of a fairly evolved basaltic magma. qasigissat–vesterdalen–nordfjord. in the area between mellemfjord and nordfjord several lava flows of basaltic andesite to andesite form the lowermost part of the nordfjord member. their total thickness is up to about 50 m (fig. 14, profiles 4–7) and several of the flows are composite. all the rocks contain xenoliths of magmamodified mudstone and sandstone, and several carry native iron. the composite flows are exposed over an area of between 150 and 200 km2 but must also have been present in the offshore areas in the west. their feeder dykes are likely to be composite similar to some of the dykes of dyke system a described below (see also fig. 174), but none have been found. an example of the diversity of the erupting pulses is illustrated by a composite andesite lava flow with native iron in the point 1300 m profile in vesterdalen (fig. 14, profile 5). the c. 12 m thick flow is composed of a lower part transitional between basaltic andesite and andesite and an upper part transitional between andesite and dacite. the lower part contains 57.0 wt% sio2, 1.7 wt% tio2, 8.1 wt% mgo and 0.57 wt% k2o. the upper part has a very inhomogeous groundmass; it contains 62.4 wt% sio2, 2.0 wt% tio2, 4.0 wt% mgo and 0.84 wt% k2o (table 6, sample 264067). here two different, both strongly contaminated magmas were emplaced during the same eruption. this flow belongs to the low-al2o3, high-p2o5 group of flows described below. low-al2o3 , high-p2o5 magmas in a group of composite lava flows. a group of one to three lava flows of andesitic to dacitic composition and sharing a unique geochemical character constitutes an important local marker horizon that was presumably produced from a single eruptive centre. the lavas are exposed along the outer part of nordfjord and extend southward to the inner part of mellemfjord, where the horizon is present at saqqarliit ilorliit (pedersen 1977b, table 7 no. 5). it belongs to the lowest part of the nordfjord member and is known from a number of profiles (fig. 14, profiles 5–8); its distribution is shown in fig. 108c. one large dacitic lava flow varies in thickness from more than 50 m around perlertut qaqqaat on the north wall of nordfjord (central disko section at 7.6 to 12.6 km) to less than 10 m. this flow forms a prominent, rusty brown horizon due to the weathering of native iron and sulfides (fig. 134). it is known from the south coast of nordfjord to mellemfjord. on the north side of nordfjord, the dacite overlies an alkali basalt lava flow and is partly covered by rhyolitic tuff from the wdgr volcano (fig. 14, profile 8; central disko section at 7.5–12.5 km). the horizon thins out just north of nordfjord and is missing in rink dal. no eruption sites have been found and there are no known intrusive rocks of similar chemistry. the thickness variations suggest that the lavas originated in the outer part of the present nordfjord. the composite mellemfjord lava flow. a large, native-ironbearing composite lava flow in the mellemfjord area on south-western disko forms an important local marker horizon in the uppermost part of the nordfjord member (figs 135, 136). the flow has been mapped photogrammetrically (pedersen 1977b, plate 1 and an isopach map, fig. 19). it extends along the northern and southern shores of the central and western parts of mellemfjord and must continue offshore well to the west of the coastline (fig. 108a). the flow must have covered an area of more than 400 km2 and its volume must have exceeded 14 km3, which makes it the most voluminous crustally contaminated lava flow on disko and among the largest lava flows within the nuussuaq basin. the flow was first discovered by k.j.v. steenstrup in 1880 and was later described by steenstrup (1883), lormf125_eqaluit_compositelava.eps fig. 133. drawing of a composite lava flow at the base of the nordfjord member near eqaluit, south-western disko. from pedersen (1977, fig. 18). 148148 enzen (1882), nicolau (1900), pauly (1969), pedersen (1977b, p. 41–46) and klöck et al. (1986). pedersen (1977b) originally assigned the flow to the niaqussat member because of its estimated high-magnesium parental composition, but it is here reassigned to the uppermost nordfjord member (see fig. 14, profile 3). the mellemfjord lava flow is lithologically very similar to the native-iron-bearing composite lava flows of the asuk member in the vaigat formation (pedersen et al. 2017), but it has a much larger volume. a simplified section through the flow is shown in fig. 137. the flow has a thin lower zone of basalt, a thin hybrid zone of basaltic andesite, and a thick upper zone of highly magnesian andesite. the lower few metres of the flow are a very fine-grained basalt without native iron (fig. 138a; table 6, sample 176565). the basalt carries very scarce plagioclase phenocrysts up to 1 mm in size and abundant pseudomorphed olivine microphenocrysts together with scarce augite microphenocrysts. the groundmass is composed of plagioclase, pigeonite and augite, ilmenite and residuum. native iron appears within the hybrid zone of basaltic andesite (table 6, sample 176556), and 2−3 m above the base the rock becomes a magnesian andesite with native iron (fig. 138b; table 6, sample 176564) and with sufficiently high cr (418 ppm) to demonstrate a highly magnesian parent. the magnesian andesite has been described by nicolau (1900) and klöck et al. (1986). it is a very finemf126_17cm1985_10_02mod_c.eps d dacite alkali basalt urdm d sed rhy tuff fig. 134. nordfjord member lava flows and rhyolitic tuff layer (rhy tuff) in the north wall of nordfjord, west disko (central disko section at around 7.8 km). the large dacite flow is 60 m thick and forms a prominent rusty brown horizon due to the weathering of native iron and sulfides. thin sediments (sed) at the base of the nordfjord member are just visible below the alkali basalt flow. urdm: upper rinks dal member. the succession is cut by a dyke (d). 149 mf127_17cmakp_1974_melfsvenokh1modcut_c.eps 530 520 urdmurdm fe ba fig. 135. the cliffs along the south coast of mellemfjord. the large native-iron-bearing composite mellemfjord lava flow (fe) forms a major marker horizon over a large area. it is the highest flow in the nordfjord member and is overlain by lava flows of the niaqussat member (unit 530). the nordfjord member (unit 520) has a basaltic andesite (ba) flow at the base. urdm: upper rinks dal member. the highest part of the cliffs reaches 900 m a.s.l. east of ivisaarqut/enok havn, west disko. mf128_17cma790213mod2_c.eps basalt andesite fig. 136. the composite mellemfjord lava flow at the north coast of mellemfjord just east of ikorfarsuit. the flow is here 75 m thick; the basaltic and andesitic parts are indicated. 150150 grained rock with an unusual assemblage of pyroxenes. there are common phenocrysts of low-ca-clinopyroxene and abundant microphenocrysts of orthopyroxene and plagioclase. the cores of some pyroxene phenocrysts show intense lamellar twinning resembling twinning in clino-hypersthene commonly found in chondrites. such twinned cores have not been found in any other andesites with native iron from disko. the magnesian andesite carries up to millimetre-sized bodies of native iron rimmed by troilite. there are scattered xenocrysts of olivine, red spinel and plagioclase with graphite, and xenoliths of magma-modified mudstone. the groundmass is composed of plagioclase, orthopyroxene and calcic clinopyroxene, ilmenite, native iron and troilite, and residual glass with a composition of potassic rhyolite. rare grains of armalcolite have been found associated with the native iron (klöck et al. 1986). in the lower part of the flow, the andesite groundmass is very fine-grained and fairly homogeneous (fig. 138b), but upwards the finegrained groundmass develops a distinct globular texture (fig. 138c) resembling a product of liquid immiscibility, similar to the texture in native-iron-bearing basaltic andesite of the asuk member. the composite mellemfjord lava flow is covered by a thin layer of lateritic claystone, which has also impregnated the uppermost metre of the scoria of the flow. the overlying lava flows are olivine-microphyric and aphyric basaltic lavas assigned to the niaqussat member. minor flows from the mellemfjord eruption site. along the north cost of mellemfjord, the composite mellemfjord lava flow is underlain by several flows packed with cognate gabbroic olivine-clinopyroxene-plagioclase nodules and magma-modified sediment inclusions up to 20 cm in size; one of the flows contains up to 4 cm long olivine grains (fig. 139). these lava flows have highly irregular upper parts where heaps of scoria alternate with thin pahoehoe lava tongues. despite the absence of a feeder body and a visible crater, these features strongly indicate close proximity to an eruption site. the first eruptions must have comprised the xenolith-rich flows, later followed by the large native-iron-bearing lava flow. composite lava flow in the kvandalen valley, eastern disko. on eastern disko, there are only two eruptive units more silicic than basalt, one in each of the nordfjord and niaqussat members. both flows are believed to be associated with the nw–se-trending dykes of system c described below (see fig. 174), which seems to have been active over a long time period. the lava flow of the nordfjord member is a large composite basaltic andesite flow with native iron in the kvandalen area. it was briefly described by pedersen & larsen (1987, fig. 2) and larsen & pedersen (1989) and shown as marker unit τ (tau) on the geological map sheet pingu (fig. 140). the lava flow extends along the north wall of kvandalen westwards from aqajaruata qaqqaa (fig. 141; fig. 10, profiles 10, 11; central disko section at 0 10 20 30 40 50 60 70 80 m 0 1 2 3 4 5 fe lateritic claystone disseminated fe fe fe red tuff fe gv04_06_035_lml fig. 137. simplified section through the composite mellemfjord lava flow; the basal part is shown expanded in the right column. the flow has a lower zone of basalt without native iron (pale red), a hybrid zone of basaltic andesite (stronger red) with disseminated native iron and three horizons of native iron accumulations (black), and a thick upper zone of highly magnesian andesite (reddish brown) with disseminated native iron (small black specks). sediment xenoliths are grey. for further description, see text. the section represents the flow along the north coast of mellemfjord between ikorfarsuit (fig. 14, profile 3) and saqqarliit silarliit (fig. 6), where the flow attains its maximum thickness of about 88 m. 151 82.5–105.4 km); its northern and eastern delimitations are unconstrained because of erosion. towards the northwest the lava flow terminates in the eastern part of the qinngusaq mountain. on the south side of kvandalen it is only present on the ridge towards charles polaris dal. the lava flow is up to 25–30 m thick and composite, with the lower few metres consisting of basaltic andesite very close to basalt (52.1 wt% sio2) while the rest of the flow is a basaltic andesite with 53–54 wt% sio2. rusty spots of weathered native iron and troilite are widespread in mf130a_8cm176565polmodcut_c.eps ol 5 mm a mf130b_8cm176564dglmodcut_c.eps fe opx pl xen 5 mm b mf130c_8cm447260aupcut1_c.eps fefe sed xen pl 5 mm c fig. 138. thin sections (scanned) of the composite mellemfjord lava flow. a: basal part of the flow: very fine-grained silicic basalt with scattered olivine (ol) that is mostly altered. no native iron is present. sample 176565, saqqarliit ilorliit, north coast of mellemfjord, western disko. b: upper zone: very fine-grained magnesian andesite with native iron (fe), orthopyroxene (opx) and a plagioclase xenocryst (pl xen). sample 176564, saqqarliit ilorliit, north coast of mellemfjord, western disko. c: upper zone: magnesian andesite with native iron (fe), a plagioclase microphenocryst (pl) and a large reacted sediment xenolith (sed xen). the fine-grained groundmass has a globular texture resembling a product of liquid immiscibility. sample 447260, jernpynten, western disko. 152152 mf131_11cmfum_1979_02_11udmod_c.eps gabbro sed ol ol fig. 139. nordfjord member lava flow with cognate gabbroic plagioclase-olivine-clinopyroxene nodules and a magma-modified sediment xenolith (sed) that is c. 8 cm long. two large olivine grains are indicated (ol). north coast of mellemfjord just east of ikorfarsuit (same locality as fig. 136). photo: finn ulff-møller. 10 km 52°30 52°15 69°45 figs 54, 55 fig. 53 fig. 156 fig. 59 fig. 141 gv01_02_157_lml ujarasussuk nuugaarsuk amisut illukunnguaq alanngoq qinngusaq illorsuaasaq qingaasaq fe ni, fe δts δs δs δt δt k is k k k fe psa pmu psa psa psa s is is is is is is is is s is is is is isis is is s is is is is is is pmu βfph2 βi βf2 βf2 βf2 βf2 βf2 βi βf1i βf1i βi βf2 βf3 βf3 βf3 βi βfph2 βfph2 βf3 βi βi βf2 1366 0.5 0.5 0.5 1 3 2.5 0.5 0.5 1.5 1 1 1 1 0.5 1.5 1 1054 1287 1.5 1.5 1.5 1 1218 870 570 1065 m moltke gletscher fre de rik la nge dal sullorsuaq • kvandalen aqajaruata qaqqaa inngigissoq qatsissup tasia charles polaris dal fig. 52 153 the upper part of the flow, which has a light rusty brown weathering colour. the flow contains scattered xenoliths of magma-modified mudstone. the flow extends for more than 23 km in a general nw–se to w–e direction and probably originally covered at least 100 km2. the erupted volume was probably in excess of 1 km3. the apparent elongation of the flow suggests emplacement from a nw–se-trending feeder dyke, which has not been located. crater site lithologies basaltic fissure eruption site on north-eastern disko. features indicative of an eruption site are seen in the second lava flow of the nordfjord member in the excellently exposed vertical coastal cliff below point 1530 m (fig. 4). the flow is ususally 15–20 m thick, but over a distance of c. 700 m its thickness increases to up to 35 m and it has a very thick scoriaceous top zone in which there are irregular lava tongues (fig. 142). the feature is interpreted as a longitudinal section through an eruption fissure. the fissure runs se–nw, parallel to the cliff face and parallel to the inferred eruption sites for the contaminated magmas (dyke system c in fig. 174). facing page: fig. 140. excerpt from the 1:100 000 scale geological map sheet 69 v.2 nord pingu, showing the extent of the native-iron-bearing lava flow (dark red) of the nordfjord member on eastern disko. other nordfjord and niaqussat member lavas are pink. the rinks dal member (rdm) flows are green (upper rdm), brown (middle rdm), and blue and olive-green (lower rdm). sediments are yellow and orange. light bluish grey areas are landslipped. the locations of some figures are indicated. annotations as on map. mf133_17cmakp_1986_06_14modcut_c.eps 520 520 fe-lava 520 518 sst fig. 141. the lava succession at the top of aqajaruata qaqqaa, eastern disko (fig. 7, profile 13). the nordfjord member (520) rests on flows of rinks dal member unit 518. fe-lava: the composite, c. 25 m thick, native-iron-bearing basaltic andesite flow shown in dark red in fig. 140. for location, see fig. 140. 154154 0 10 0 20 0 30 0 40 0 50 0 60 0 70 0 80 0 90 0 10 00 m n iaq us sa t m em be r n or df jo rd m em be r tw o su cc es sio ns o f l av a flo w s s ep ar at ed b y re d se di m en t tw o su cc es sio ns o f l av a flo w s s ep ar at ed b y re d se di m en t la va fl ow s s lig ht ly yo un ge r t ha n th e er up tiv e fis su re ye llo w ish tu ffa ce ou s p yr oc las tic ro ck s py ro cla st ic de po sit s f ro m e ru pt ive fi ss ur e m as siv e ba sa lt lav a flo w fr om e ru pt ive fi ss ur e n or df jo rd m em be r lo w es t l av a flo w in n or df jo rd m em be r w hi te rh yo lit ic tu ff, re de po sit ed th re e lav a flo w s sc or iac eo us la va to p c ol um na r j oi nt in g ri nk s d al m em be r 15 00 se n w m 14 00 13 00 14 50 13 50 g v0 3_ 03 _0 34 _l m l fi g. 1 42 . l on gi tu di na l s ec tio n th ro ug h a b as al tic fi ss ur e e ru pt io n sit e i n th e n or df jo rd m em be r. t he fi ss ur e r un s s e– n w , p ar al le l t o th e c lif f f ac e. ph ot og ra m m et ric al ly m ea su re d se ct io n al on g t he co as ta l c lif f be lo w p oi nt 15 30 m , n or th -e as te rn d isk o. f or lo ca tio n, se e f ig s 4 , 6 . 155 craters and lava flows around point 440 m near hammer dal. the top lava in the point 600 m profile north of hammer dal is a more than 45 m thick, native-iron-bearing dacite with disseminated graphite (figs 104c, 132) which was formerly assigned to the niaqussat member (pedersen 1977a, fig. 7), but which is here reassigned to the nordfjord member (fig. 14, profile 9). the lava was erupted from a crater site c. 1.7 km west-north-west of point 600 m exposed in a gully close to point 440 m and described as the point 440 crater by pedersen (1977a, p. 11–13, fig. 9). the locality shows preserved parts of a more than 250 m wide crater with a height of at least 110 m which is easily distinguished by rocks with a yellowish-grey colour speckled with red (fig. 143; see also the photogrammetric interpretation in fig. 104b). the crater is composed of a mushroom-shaped central feeder body of massive, native-iron-bearing dacite surrounded by vesiculated and fragmented, oxidised dacite forming a crater breccia rich in dark grey, graphite-bearing magmamodified mudstone xenoliths (fig. 144). the crater deposit is covered by a thin layer of lateritic claystone. the crater was subsequently overflowed and gradually covered by more than six basaltic lava flows of the niaqussat member. these flows are olivine-microphyric pahoehoe lavas; the lowermost flow is 45 m thick and formed a local lava lake when it filled part of the crater. the succession was later tilted to dip 24°w. mf135ab_17cma1974_k440krvestråfil_c.eps 530 la dip 24°w 530 la 530 la cr feeder cr d cr fig. 143. exposed section through the point 440 crater north of hammer dal, north-western disko. the light rock (cr) is a native-iron-graphite dacite crater breccia with a feeder body; the dark rocks comprise later infill of lava flows of the niaqussat member (530 la). the succession dips 24°w and is cut by a later dyke (d). the red circle indicates the position of fig. 144. see text for detailed descriptions. for locality, see fig. 103, loc. 7. mf136_8cma1974_k440kraterbreccnærmodcut_c.epsfig. 144. the volcanic breccia in the point 440 crater (see fig. 143). the breccia is composed of vesiculated and fragmented, oxidised dacite that is rich in dark grey, graphite-bearing, magma-modified mudstone xenoliths. length of hammer 47 cm. 156156 the site illustrates how an area rich in local craters was re-shaped into a lava plateau after the onset of the regional basaltic volcanism of the niaqussat member. craters and lava flows between hanekammen, hammer dal and giesecke dal. the upper part of the nordfjord member (from the top of the dacites described above) in an area extending from hanekammen and hammer dal to giesecke dal, north-western disko, is dominated by basaltic to andesitic lavas and craters decribed by pedersen (1977a, figs 1 and 2). the western part of the area has been extensively faulted, subsequently eroded into a tertiary planation surface that now has a westerly dip (bonow et al. 2006), cut by large incised valleys, and eroded again by the holocene sea. it is now mostly covered either by quarternary deposits or by extensive low-lying marine deposits. despite the generally poor exposure, excellent sections are exposed in the north wall of hammer dal around point 600 m (fig. 14, profile 9; fig. 104c), in two e–w-trending gullies north of hammer dal (fig. 14, profiles 10, 11; fig. 104a, b), and in the mountain walls around point 500 m south of niaqussat (fig. 14, profile 12). the area contains scattered craters and volcanic necks, and many more must exist undetected beneath the quaternary cover. some of the craters and necks are connected to feeder dykes such as the hammer dal complex (ulff-møller 1977) and the hanekammen complex (ulff-møller 1990). some of the lavas and craters were probably fed by dyke system a described below. the basalts and andesites of the upper part of the nordfjord member are extensively sediment-contaminated and were clearly fed from high-level magma reservoirs; many lavas and crater breccias are packed with magmamodified mudstones and sandstones as well as cognate igneous inclusions (autoliths). however, the volcanism at the individual sites was short-lived, and except for the elusive west disko graphite rhyolite volcano there is no evidence of large central volcanoes that sustained longlived activity. tunup qaqqaa, south-western disko. the mountaintops in the area around tunup qaqqaa north of the inner part of kangerluk are capped by lavas of the nordfjord and niaqussat members (pedersen & ulff-møller 1987; south disko section at 35.5–42 km). a feeder crater site is exposed in the mountain wall facing kangerluk (fig. 87). the crater was part of the feeder system for a c. 40 m thick basaltic andesite lava erupted at the base of the nordfjord member (fig. 8, profile 1, sample 175077). the lava flow probably covered more than 20 km2 with a volume of around 0.5–1 km3; it is likely to be composite, but present knowledge is incomplete. chemical compositions of the nordfjord member the nordfjord member comprises an almost continuous compositional suite from basalt over basaltic andesite, andesite and dacite to rhyolite. it should, however, be noted that these rocks are distinctly different from orogenic rocks with the same names. in particular, the nordfjord member rocks have much higher contents of mgo, ni, and cr than their orogenic namesakes. table 6 shows representative chemical analyses of the nordfjord member rocks, and all analyses are plotted in the variation diagrams in figs 145–147. in contrast to the variation diagrams for the rinks dal member, mgo was chosen as the main variation parameter because the nordfjord member rocks show strongly divergent trends in feo* that invalidate the mg-number as a common variation parameter. major elements the nordfjord member rocks are subalkaline, with total alkalis for basalts and basaltic andesites less than 3.5 wt% except for the above mentioned single alkali basalt flow which has 4 wt% total alkalis. many basalts are quartz normative. of the analysed samples four andesites, three dacites, and all the rhyolites are corundum normative. the alkali basalt samples are either slightly nepheline or slightly hypersthene normative, demonstrating its mildly alkaline character. the nordfjord member basalts overlap compositionally with the rinks dal member basalts but with a displacement towards slightly higher sio2 and lower cao, tio2 and na2o contents (fig. 145). the nordfjord member basalts have on average 49.8 wt% sio2, 31% of the analyses have sio2 >50 wt%, and 76% are quartz normative. in comparison, the rinks dal member basalts have on average 49.2 wt% sio2, 6% of the analyses have sio2 >50 wt%, and 35% are quartz normative. feo* contents are similar to those in the rinks dal member. the single alkali basalt flow (four analysed samples) is particularly discernible by its high na2o. it has correspondingly low cao and low sio2, and high tio2, k2o and p2o5. basaltic andesites and andesites span a larger range in mgo than the basalts and form trends parallel to the basalts at higher sio2 and k2o and lower feo*, cao, 157 45 50 55 60 65 70 75 80 0 2 4 6 8 10 rhyolite dacite mellemfjord lava flow basaltic andesite and andesite basalt alkali basalt flow rinks dal mb 5 10 15 20 0 2 4 6 8 10 rinks dal mb 0 5 10 15 0 2 4 6 8 10 rinks dal mb 0 5 10 15 0 2 4 6 8 10 rinks dal mb 0 1 2 3 4 5 0 2 4 6 8 10 rinks dal mb 0 1 2 3 4 5 0 2 4 6 8 10 rinks dal mb 0 1 2 3 4 5 6 0 2 4 6 8 10 mgo (wt%) rinks dal mb 0 2 4 6 8 10 mgo (wt%) rinks dal mb sio2 al2o3 feo* cao tio2 na2o k2o p2o5 0.8 0.6 0.4 0.2 0.0 fig. 145. major-element variation diagrams for rocks of the nordfjord member. data in wt% oxides recalculated volatile-free. feo* is total iron as feo. dotted outlines in the al2o3, p2o5 and k2o diagrams show the fields of the low-al2o3, high-p2o5 andesites and dacites. 158158 0 50 100 150 200 0 2 4 6 8 10 rhyolite dacite mellemfjord lava flow basaltic andesite and andesite basalt alkali basalt flow 0 100 200 300 400 500 600 700 800 900 0 2 4 6 8 10 0 100 200 300 400 500 600 700 800 900 0 2 4 6 8 10 0 10 20 30 40 50 60 70 0 2 4 6 8 10 0 100 200 300 400 0 2 4 6 8 10 0 5 10 15 20 25 30 35 0 2 4 6 8 10 0 2 4 6 8 10 12 14 0 2 4 6 8 10 0 20 40 60 80 100 120 0 2 4 6 8 10 rinks dal mb rinks dal mb rinks dal mb rinks dal mb rinks dal mb rinks dal mb rinks dal mb rinks dal mb rb ba sr y mgo (wt%) zr mgo (wt%) nb zr/y ce fig. 146. incompatible trace element variation diagrams for rocks of the nordfjord member. data in ppm. dotted outlines in some diagrams show the fields of the low-al2o3, high-p2o5 andesites and dacites. 159 tio2 and na2o. in contrast to the basalts, feo* is constant or decreases with decreasing mgo. the dacites continue the trend from the andesites towards higher sio2 and k2o and lower feo* and cao and show no increase in tio2 and na2o. a subgroup of andesites and dacites form separate trends, which are contoured in some of the diagrams in fig. 145. these rocks have particularly high p2o5 and also high sio2, low al2o3, and low k2o. they represent a group of related lava flows in a limited area of about 25 km2 around nordfjord and southwards to mellemfjord. these flows are regarded as predominantly sandstonecontaminated, as discussed later. the rhyolites contain 70–77 wt% sio2 and 0.2–1.2 wt% mgo and are the most evolved rocks on disko. they comprise two compositional subgroups (best seen in the k2o, tio2 and cao diagrams, fig. 145), a lowsilica group (garnet rhyolites) with lower sio2 and k2o rhyolite dacite mellemfjord lava flow basaltic andesite and andesite basalt alkali basalt flow 0 100 200 300 400 500 600 0 2 4 6 8 10 0 2 4 6 8 10 0 10 20 30 40 50 0 2 4 6 8 10 0 100 200 300 400 500 0 50 100 150 200 0 2 4 6 8 10 0 100 200 300 400 0 2 4 6 8 10 0 200 400 600 800 1000 1200 0 2 4 6 8 10 rinks dal mb rinks dal mb rinks dal mb rinks dal mb rinks dal mb rinks dal mb v sc cu zn mgo (wt%) ni mgo (wt%) cr fig. 147. transition-element variation diagrams for rocks of the nordfjord member. data in ppm. dotted outline in the v diagram shows the field of the low-al2o3, high-p2o5 andesites and dacites. 160160 and higher al2o3, cao, feo*, tio2 and p2o5 than the other, high-silica group (sanidine rhyolites). the rocks contain 0.9–1.2 wt% feo and the glasses have very low mgo/(mgo+feo) ratios of 0.04–0.08, indicative of a reduced state. the most unusual feature of the strongly contaminated volcanic rocks of the nordfjord member is the presence of native iron (iron-carbon alloys), sulfur and carbon in many of these rocks; analytical data for these elements are given in table 7. the highest content of native iron analysed is 2.4 wt% fe0 in an andesite, and the highest toc is 3.1 wt% in a rhyolite (apart from a sediment xenolith with 11.72 wt% toc). the rhyolites do not contain native iron but are rich in graphite. sulfur in rhyolite is generally low and has disappeared by degassing of the flow during eruption. 1 10 100 1000 la ce pr nd sm eu gd tb dy ho er tm yb lu basalts, sio2 = 48–50 wt% alkali basalt 1 10 100 1000 la ce pr nd sm eu gd tb dy ho er tm yb lu highest sio2 basalts, sio2 = 50–52 wt% 1 10 100 1000 rb ba th u nb ta k la ce pb pr sr nd p sm zr hf eu ti gd tb dy y yb 1 10 100 1000 rb ba th u nb ta k la ce pb pr sr nd p sm zr hf eu ti gd tb dy y yb 1 10 100 1000 la ce pr nd sm eu gd tb dy ho er tm yb lu basaltic andesites, sio2 = 52–57 wt% andesites, normal p2o5 1 10 100 1000 rb ba th u nb ta k la ce pb pr sr nd p sm zr hf eu ti gd tb dy y yb 1 10 100 1000 la ce pr nd sm eu gd tb dy ho er tm yb lu 1 10 100 1000 rb ba th u nb ta k la ce pb pr sr nd p sm zr hf eu ti gd tb dy y yb fig. 148 (first of two parts). ree and multi-element diagrams for representative rocks of the nordfjord member. dacites with ‘normal p2o5’ comprise high-ni, high-al and iron-ilmenite dacites. left diagram, chondrite normalised; right diagram, primitive mantle normalised; normalisation factors from mcdonough & sun (1995). 161 trace elements variation diagrams for trace elements are shown in figs 146, 147. the basalts have trace-element contents within the same range as the rinks dal member basalts although high values of sr, y, zr, nb, and cu are not attained. the more silicic rocks have higher rb, ba, ce and cr, and lower v and cu. apparent trends for sr, y, zr, nb, sc and zn pass through the level of the basalts but do not increase, or even decrease, with decreasing mgo. ni shows rather erratic variations, which are tied to the presence or absence of native iron or sulfides in some flows. the high-p2o5 andesites and dacites have high zr, nb and zr/y and distinctly low v without having low tio2. the high-silica sanidine rhyolites have higher rb, lower sr and v, and particularly much lower ba, zr and ce than the low-silica garnet rhyolites. 1 10 100 1000 la ce pr nd sm eu gd tb dy ho er tm yb lu 1 10 100 1000 rb ba th u nb ta k la ce pb pr sr nd p sm zr hf eu ti gd tb dy y yb 1 10 100 1000 la ce pr nd sm eu gd tb dy ho er tm yb lu 0 1 10 100 1000 rb ba th u nb ta k la ce pb pr sr nd p sm zr hf eu ti gd tb dy y yb 1 10 100 1000 la ce pr nd sm eu gd tb dy ho er tm yb lu mellemfjord composite lava flow andesite basaltic andesite basalt garnet rhyolites sanidine rhyolites 1 10 100 1000 rb ba th u nb ta k la ce pb pr sr nd p sm zr hf eu ti gd tb dy y yb 1 10 100 1000 la ce pr nd sm eu gd tb dy ho er tm yb lu 1 10 100 1000 rb ba th u nb ta k la ce pb pr sr nd p sm zr hf eu ti gd tb dy y yb dacites, normal p2o5 andesites and dacites, low al2o3, high p2o5 dacite andesite fig. 148 (second of two parts). ree and multi-element diagrams for representative rocks of the nordfjord member. dacites with ‘normal p2o5’ comprise high-ni, high-al and iron-ilmenite dacites. left diagram, chondrite normalised; right diagram, primitive mantle normalised; normalisation factors from mcdonough & sun (1995). 162162 the alkali basalt has high contents of incompatible trace elements, particularly sr and nb, and low sc and cu. ree and multi-element diagrams are shown in fig. 148. the ree patterns of the basalts are fairly similar to those of the basalts of the rinks dal member except for slightly increased la–nd limbs in five out of nine samples; these do not possess the humpback-shaped characteristic of the rinks dal member. the basaltic andesites have significantly increased la–nd and almost straight ree patterns; a few have negative eu anomalies. the andesites and dacites have ree patterns parallel displaced towards higher values than the basaltic andesites, and, apart from the high-p2o5 group, have distinct negative eu anomalies. the high-p2o5 group has ree patterns without eu anomalies and lower hree than the other rocks. the low-silica garnet rhyolites have very high ree patterns nearly parallel to those of the low-p2o5 dacites, whereas the high-silica sanidine rhyolites have very different patterns with decreased lree and very deep eu anomalies. the multi-element patterns show that many of the basalts have slightly increased ba–th–u relative to the rinks dal member. basalts with <50 wt% sio2 have diminished k and pb troughs, whereas basalts with >50 wt% sio2 have variable k and no pb troughs. a high-tio2 basalt (328463, table 6) has distinct peaks at ti and p while other incompatible elements are not increased; there are four analyses (probably representing three flows, all on western disko) of this basalt type (fig. 145) which seems to have acquired its high ti and p by other means than normal fractionation. the basaltic andesites have further increased rb–ba–th–u limbs, nb–ta troughs, k and pb peaks and small zr–hf peaks. the normal-p2o5 andesites and dacites have deepened nb–ta troughs because the elements on both sides are increased more than nb and ta; they also have distinct k and pb peaks and relative troughs at sr, p, eu and ti. the high-p2o5 andesites and dacites have lower contents of many incompatible elements but higher p and zr–hf– eu than the corresponding low-p2o5 rocks. the garnet rhyolites have the highest contents of incompatible elements; a deep ti trough indicates ti-oxide fractionation. the sanidine rhyolites appear to have fractionated feldspar (very low ba, sr, eu), apatite, ti-oxide, zircon, and monazite. monazite, which is identified petrographically, has removed th (but not u) and light to middle ree. composite lava flows the nordfjord member contains a number of composite lava flows as described above and in pedersen (1977b). the chemical analyses (fig. 149) show that the variation within the composite flows mainly occurs in the basaltic andesite compositional interval (52–57 wt% sio2). the basal parts of the flows are generally more mgo-rich and sio2-poor than their main parts, with a few exceptions where the opposite is the case. other elements behave variably and may decrease from base to main part in some flows and increase in others (e.g. tio2 and cao), testifying that the variation is not caused by ordinary crystal fractionation processes. the incompatible elements normally increase from the base to the main part of a flow (na2o, p2o5 and many trace elements). most of the compatible trace elements, including v, decrease from base to main part, but ni increases. the composite mellemfjord lava flow has a compositional range of 50.9–61.1 wt% sio2, 8.8–5.7 wt% mgo, 10.9–6.6 wt% cao, 0.36–1.78 wt% k2o, and 12.5–6.8 wt% feo*. the least contaminated basalt sample has 10.2 wt% feo*; a single sample with higher feo* than this value (12.5 wt% feo*) has accumulated native iron. 163 eqaluit mellemfjord lava vesterdalen 1012 vesterdalen 1300 nordfjord 1137 kvandalen cr sio2 cao tio2 sr zr v ni 0 200 400 600 800 1000 1200 0 2 4 6 8 10 mgo (wt%) 45 50 55 60 65 0 2 4 6 8 10 5 10 15 0 2 4 6 8 10 1 2 3 0 2 4 6 8 10 0 100 200 300 400 0 2 4 6 8 10 0 100 200 300 0 2 4 6 8 10 0 100 200 300 400 0 2 4 6 8 10 0 50 100 150 200 250 300 350 0 2 4 6 8 10 mgo (wt%) fig. 149. variation diagrams for composite lava flows of the nordfjord member. 5201 5201 5201 5201 5201 5201 5201 5201 5201 5201 5201 5201 156678 327011 332882 279298 326647 362398 279465 328463 318768 279297 263934 318809 5440.66 5322.36 5221.55 5431.57 5438.32 5240.68 5433.37 5359.31 5151.96 5431.57 5439.11 5210.29 7002.622 6917.510 7028.983 6955.205 7015.636 6950.968 7000.524 6923.916 7006.107 6955.189 6954.400 6944.761 406 921 1679 951 377 1267 1092 818 1238 931 761 1036 45.82 49.06 48.18 50.77 49.57 48.87 48.91 47.04 49.12 48.80 54.22 53.04 3.49 1.65 1.49 1.55 1.98 2.23 2.60 3.64 2.90 1.39 1.55 1.72 15.91 14.35 14.82 16.06 14.79 15.20 13.40 13.42 12.94 13.25 13.46 15.03 4.27 6.38 3.85 1.49 3.25 2.98 4.23 6.49 4.60 6.67 2.25 1.04 8.00 4.93 6.90 8.20 7.91 8.43 9.70 7.77 11.03 3.52 7.37 8.12 0.17 0.18 0.17 0.16 0.17 0.17 0.22 0.20 0.23 0.14 0.15 0.17 6.90 8.69 8.40 8.11 7.53 6.36 6.20 6.14 5.24 9.30 9.05 7.72 8.46 10.59 11.79 9.96 10.91 11.23 10.43 10.43 9.95 7.97 7.94 9.12 3.28 2.02 2.01 1.97 2.05 2.30 2.46 2.43 2.39 1.51 1.38 2.17 0.630 0.307 0.136 0.189 0.250 0.133 0.260 0.244 0.300 0.842 0.660 0.622 0.390 0.154 0.122 0.178 0.192 0.220 0.254 0.368 0.286 0.146 0.230 0.198 1.92 1.70 1.48 1.85 2.02 1.06 1.51 1.53 1.05 5.32 1.17 0.88 99.24 100.01 99.35 100.49 100.62 99.18 100.17 99.69 100.04 98.86 99.43 99.83 11.84 10.67 10.36 9.54 10.83 11.11 13.51 13.61 15.17 9.52 9.39 9.06 54.10 62.22 62.11 63.23 58.43 53.66 48.15 47.71 41.13 66.39 66.08 63.29 92.0 93.1 77.3 82.3 94.1 96.2 116 106 128 80.1 76.8 86.2 56.2 107 154 21.8 111 131 173 318 255 26.5 66.9 71.3 76.6 136 142 77.4 116 73.7 63.4 84.8 47.0 39.0 88.2 165 21.0 32.0 38.4 28.8 38.9 36.4 35.3 32.0 36.2 29.4 27.7 32.1 311 282 317 233 319 334 398 360 432 265 206 261 94.1 562 407 630 411 233 95.4 139 26.5 1198 830 552 18.4 19.3 18.3 20.5 19.8 21.9 23.2 23.9 24.2 18.1 19.2 21.4 10.5 5.26 0.83 2.02 1.51 1.09 3.89 1.51 4.71 21.0 20.4 19.3 574 195 186 246 183 224 248 230 243 177 203 226 30.2 24.6 24.1 23.7 30.6 32.5 36.4 34.1 38.4 18.2 23.0 29.0 213 99.9 81.1 141 130 150 155 138 178 110 188 161 17.5 4.41 5.12 6.37 5.56 6.48 7.72 6.48 8.12 4.88 7.03 7.09 0.137 0.087 0.004 0.065 0.007 0.022 0.050 0.116 0.072 0.543 0.788 0.477 124 56.0 44.9 130 63.0 52.2 55.7 37.8 50.1 137 224 181 13.9 5.79 6.13 12.8 7.36 8.27 8.85 7.48 8.73 8.53 13.2 13.5 36.5 15.3 14.4 28.9 19.4 22.4 23.6 19.9 24.9 20.5 29.5 32.1 5.21 2.20 2.08 3.86 2.91 3.43 3.86 3.38 4.02 2.89 4.05 4.28 25.2 11.5 10.3 16.7 14.6 17.5 19.4 17.2 21.2 13.0 17.3 19.4 6.61 3.52 3.29 4.01 4.44 5.24 5.66 5.14 6.36 3.53 4.29 4.97 2.31 1.31 1.20 1.36 1.53 1.74 1.95 1.79 2.21 1.19 1.42 1.51 7.06 4.22 4.12 4.28 5.44 6.15 6.43 5.83 7.67 4.00 4.39 5.78 1.05 0.713 0.705 0.690 0.885 1.01 1.06 0.984 1.243 0.590 0.727 0.882 5.89 4.22 4.24 4.07 5.24 6.01 6.27 5.82 7.27 3.44 4.10 5.17 1.13 0.877 0.874 0.807 1.094 1.20 1.23 1.17 1.42 0.676 0.880 1.06 3.01 2.36 2.41 2.09 3.00 3.10 3.22 3.04 3.77 1.77 2.16 2.79 0.416 0.330 0.351 0.308 0.428 0.453 0.459 0.449 0.548 0.244 0.313 0.382 2.45 2.05 2.07 1.91 2.58 2.73 2.79 2.66 3.19 1.53 2.00 2.48 0.343 0.285 0.302 0.280 0.367 0.393 0.410 0.371 0.454 0.224 0.284 0.347 5.14 2.59 2.15 3.60 3.35 3.80 4.06 3.62 4.61 2.82 4.57 4.01 1.15 0.285 0.289 0.391 0.344 0.386 0.518 0.405 0.569 0.290 0.470 0.679 1.66 1.32 0.583 3.28 1.69 1.58 1.22 1.00 1.63 2.74 3.37 3.17 1.09 0.669 0.394 2.50 0.930 0.977 0.731 0.592 0.665 1.44 2.45 2.51 0.355 0.198 0.129 0.429 0.292 0.260 0.214 0.155 0.223 0.375 0.645 0.641 0.703419 0.703303 0.707465 0.708146 -16.06 -17.70 41.37 51.04 0.512964 0.512883 0.512286 0.512222 7.87 6.28 -5.36 -6.62 17.910 18.430 16.936 16.978 15.388 15.482 15.045 15.079 37.677 38.172 37.225 37.219 lith. code ggu no. deg. w deg. n altitude, m sio2 tio2 al2o3 fe2o3 feo mno mgo cao na2o k2o p2o5 volatiles sum feo* mg number zn cu ni sc v cr ga rb sr y zr nb cs ba la ce pr nd sm eu gd tb dy ho er tm yb lu hf ta pb th u isotope ratios calculated at 60 ma 87sr/86sr60 εsr 143nd/144nd60 εnd 206pb/204pb60 207pb/204pb60 208pb/204pb60 for explanation of lithological codes, see table 1. for petrographical notes on the samples, see table 6d–e. geographical coordinates in wgs 84. first two digits are degrees, then follow minutes in decimal form. major elements in wt% (xrf analyses). trace elements in ppm (zn–ga: xrf analyses; rb–u: icp-ms analyses). feo* = total iron as feo. mg number = 100 × atomic mg/(mg+fe2+), with the iron oxidation ratio adjusted to fe2o3/feo = 0.15. table 6a. chemical analyses of rocks of the nordfjord member basalt basaltic andesitelithology alkali basalt 165 5201 5201 5201 5201 5201 5201 5201 5201 5201 5201 5201 5201 264087 176626 274430 176565 176556 176564 176579 176582 176448 176473 113463 176411 5432.80 5452.49 5441.97 5447.61 5432.14 5447.61 5444.68 5444.68 5439.89 5441.43 5438.72 5438.76 6949.408 6944.605 7010.578 6947.511 6944.859 6947.511 6938.688 6938.688 7009.282 7012.046 7015.447 7015.246 802 128 309 456 821 459 245 241 500 360 410 368 56.11 52.55 56.19 49.90 54.60 59.93 52.93 54.87 57.40 58.71 56.93 59.80 1.42 2.17 1.71 1.62 1.18 1.11 2.18 3.06 1.75 1.72 1.86 1.99 15.14 15.17 15.34 14.33 13.19 13.49 15.11 14.14 14.55 14.75 16.14 15.57 2.69 1.84 1.06 2.02 4.13 0.62 2.02 1.33 0.95 2.54 0.00 0.00 5.74 8.40 7.57 8.21 8.76 6.83 8.15 9.84 7.27 4.77 10.85 8.48 0.12 0.19 0.15 0.20 0.16 0.15 0.17 0.17 0.15 0.12 0.13 0.14 6.20 6.06 5.28 8.59 6.94 5.76 5.81 3.57 5.25 4.44 3.37 2.48 8.16 9.72 7.64 10.72 7.69 6.70 9.50 8.35 7.42 6.32 6.75 5.79 1.91 2.52 2.10 2.06 2.26 2.41 2.54 2.45 2.27 2.34 1.89 2.27 1.210 0.510 1.313 0.420 1.120 1.690 0.930 0.880 1.320 1.490 1.347 1.640 0.170 0.240 0.237 0.130 0.165 0.150 0.230 0.290 0.190 0.270 0.273 0.340 0.82 0.79 1.93 1.55 0.66 1.20 1.07 1.14 1.32 1.65 0.95 1.16 99.69 100.16 100.52 99.75 100.86 100.04 100.64 100.09 99.94 99.12 100.49 99.66 8.16 10.06 8.52 10.03 12.48 7.39 9.97 11.04 8.12 7.06 10.85 8.48 60.58 54.93 55.61 63.41 52.94 61.19 54.11 39.55 56.65 56.00 38.58 37.17 0.09 0.22 0.03 0.18 0.00 0.41 81.2 89.3 86.0 80.3 62.9 55.7 92.6 106.3 90.2 89.6 80.4 81.2 12.7 11.3 71.9 225 78.4 79.9 11.7 28.4 68.1 63.4 110 101 14.8 10.3 67.5 84.1 85.8 113 9.1 65.0 63.6 69.2 87.2 90.4 28.5 30.4 31.0 33.3 29.7 22.3 33.7 34.0 29.1 28.8 28.3 26.0 238 211 215 272 218 160 236 206 218 209 186 155 442 359 301 586 507 411 365 128 334 297 152 106 20.5 21.8 22.1 18.7 19.4 19.1 22.5 23.9 21.3 21.6 23.7 23.3 34.0 18.8 33.7 9.07 38.2 46.8 29.7 29.7 35.6 43.2 40.2 50.3 240 268 208 189 178 172 255 266 197 190 239 234 23.3 26.6 33.6 26.6 30.7 31.6 27.6 30.9 35.0 35.2 35.3 34.7 145 185 222 104 180 214 182 258 230 238 231 262 7.29 8.19 9.08 4.42 6.44 7.39 8.09 10.8 8.84 9.51 10.8 11.5 0.499 0.408 0.895 0.123 0.847 0.742 0.875 1.03 0.932 1.03 0.993 1.28 251 178 284 79.78 291 358 193 271 285 322 364 398 15.7 13.9 19.8 6.88 17.2 21.2 14.3 18.8 19.9 21.1 24.8 26.5 33.3 32.6 45.7 18.1 37.6 47.8 32.5 41.2 45.7 48.9 53.1 59.1 4.51 4.24 5.85 2.68 4.93 6.05 4.57 5.89 5.72 6.17 7.01 7.19 18.6 19.2 25.2 13.3 20.9 25.5 20.3 25.0 25.1 26.7 29.3 30.1 4.25 4.93 6.30 4.08 5.13 6.05 5.08 5.95 6.32 6.57 6.49 6.95 1.25 1.67 1.57 1.30 1.20 1.17 1.66 1.91 1.55 1.54 1.62 1.64 4.27 5.65 7.01 4.92 5.32 6.71 5.17 6.10 7.20 7.10 6.72 7.66 0.693 0.847 1.056 0.804 0.898 1.033 0.885 0.989 1.096 1.113 1.073 1.111 3.95 4.85 6.10 4.74 5.24 5.84 5.09 5.77 6.34 6.39 6.06 6.33 0.782 0.968 1.24 0.965 1.09 1.17 1.05 1.14 1.26 1.24 1.18 1.25 2.10 2.62 3.31 2.61 2.80 3.13 2.66 2.93 3.50 3.47 3.24 3.42 0.302 0.357 0.483 0.373 0.419 0.451 0.380 0.406 0.488 0.504 0.479 0.497 1.89 2.29 2.93 2.17 2.58 2.69 2.24 2.57 3.08 3.01 2.90 2.88 0.280 0.315 0.417 0.311 0.361 0.386 0.335 0.378 0.441 0.439 0.437 0.430 3.76 4.59 5.63 2.81 4.70 5.42 4.78 6.33 5.98 6.17 5.75 6.61 0.478 0.561 0.561 0.311 1.080 0.963 0.544 0.935 0.818 0.623 0.699 0.817 5.43 3.68 5.94 1.83 4.27 5.14 3.98 4.78 6.16 7.01 6.82 6.93 3.30 2.36 4.27 1.03 4.02 5.19 2.47 3.41 4.22 4.62 5.64 5.64 0.855 0.621 1.10 0.286 1.04 1.37 0.673 0.901 1.13 1.22 1.33 1.43 0.708947 0.704900 0.709988 0.709660 0.714259 62.40 4.96 77.17 72.53 137.79 0.512195 0.512595 0.511918 0.511766 -7.14 0.66 -12.54 -15.50 16.892 17.072 16.763 16.907 15.044 15.105 15.010 15.041 37.177 37.306 37.112 37.172 lith. code ggu no. deg. w deg. n altitude, m sio2 tio2 al2o3 fe2o3 feo mno mgo cao na2o k2o p2o5 volatiles sum feo* mg number toc wt% s wt% zn cu ni sc v cr ga rb sr y zr nb cs ba la ce pr nd sm eu gd tb dy ho er tm yb lu hf ta pb th u isotope ratios calculated at 60 ma 87sr/86sr60 εsr 143nd/144nd60 εnd 206pb/204pb60 207pb/204pb60 208pb/204pb60 andesitebasaltic andesitelithology table 6b. chemical analyses of rocks of the nordfjord member mellemfjord composite lava flow eqaluit composite flow 166166 5201 5201 5201 5201 5201 5201 5201 5207 5207 5207 5207 5207 264064 264067 176441 176466 176471 176555 176443 326550 326465 156518.1 156516 156559 5428.35 5428.36 5439.52 5450.27 5449.50 5432.29 5440.02 5439.65 5446.30 5446.31 5446.31 5444.46 6952.312 6952.302 7009.224 7011.902 7012.836 6944.730 7009.311 7010.035 7007.630 7007.633 7007.633 7004.514 998 990 434 3 1 734 525 417 164 160 160 266 59.67 61.62 62.87 63.86 66.20 62.86 64.07 59.15 70.05 70.24 72.49 73.35 1.93 2.01 1.74 1.62 1.25 2.27 1.25 1.64 0.48 0.47 0.09 0.04 13.05 12.86 14.46 14.34 13.59 12.82 14.65 15.80 13.60 12.88 12.39 12.63 1.62 1.07 1.20 0.00 2.22 1.05 0.00 0.00 0.84 1.16 0.72 0.57 6.72 6.82 6.84 7.61 4.39 6.88 6.52 4.93 2.55 2.78 0.68 0.60 0.15 0.14 0.12 0.13 0.11 0.16 0.10 0.10 0.06 0.05 0.03 0.03 5.19 3.96 2.02 1.49 1.29 2.03 2.05 2.81 0.64 0.54 0.16 0.17 7.19 6.78 5.47 4.98 3.62 6.34 3.63 5.21 2.05 2.12 0.58 0.58 2.10 2.15 2.43 2.48 2.67 2.45 2.28 1.88 2.40 2.50 1.76 2.68 0.610 0.830 1.830 1.990 2.340 0.850 2.150 1.358 3.848 3.820 5.730 5.100 0.420 0.460 0.290 0.300 0.310 0.500 0.250 0.181 0.198 0.200 0.070 0.090 1.06 1.12 0.98 0.95 1.40 1.69 2.40 6.19 3.33 2.61 4.15 3.59 99.71 99.82 100.25 99.75 99.39 99.90 99.35 99.24 100.05 99.37 98.85 99.43 8.18 7.78 7.92 7.61 6.39 7.82 6.52 4.93 3.31 3.82 1.33 1.11 56.21 50.72 34.03 28.37 29.00 34.42 38.87 53.53 28.14 22.22 19.60 23.60 0.07 0.27 1.17 2.72 0.17 0.22 0.01 0.19 0.22 0.28 0.64 0.02 85.1 88.5 80.67 89.0 79.7 92.9 68.2 89.6 62.1 57.1 60.7 57.6 29.5 21.7 47.71 41.1 58.6 31.2 131 86.7 7.50 9.13 4.10 7.67 44.0 20.1 63.91 71.5 59.9 44.1 201 28.8 5.57 4.86 2.09 1.33 32.1 31.8 25.37 22.9 20.5 32.6 20.2 28.5 13.2 17.8 8.73 6.23 144 116 96.20 70.5 60.5 89.5 108 164 20.9 21.6 5.72 0.80 389 244 62.64 36.8 37.1 64.5 124 145 9.19 6.96 0.80 1.29 20.7 21.1 22.88 22.3 22.3 21.7 20.6 26.3 22.8 22.1 22.7 22.3 19.7 27.8 48.50 56.1 65.9 26.9 61.1 52.7 121 125 197 211 264 274 242.37 249 193 304 198 225 174 172 29 18 29.8 31.1 33.30 33.5 42.2 31.7 29.2 37.1 37.8 40.3 53.0 25.0 320 355 252.17 275 342 436 267 240 69.3 83.4 60.5 37.4 11.2 12.2 10.92 11.5 12.9 13.9 11.3 11.1 12.3 12.1 9.82 9.73 0.537 0.688 0.793 1.06 1.11 0.635 1.41 1.49 2.86 2.94 5.69 7.12 322 295 473.81 485 536 346 491 365 798 769 181 66.2 20.2 22.1 27.23 28.7 32.6 24.0 30.5 25.6 42.2 43.3 14.9 8.79 44.1 47.6 57.13 62.9 72.5 53.7 66.0 51.4 92.3 92.7 36.3 22.1 5.96 6.33 7.50 7.60 9.13 7.09 7.78 7.31 10.97 11.44 4.81 2.89 25.4 27.1 30.23 31.2 38.0 31.1 31.0 29.0 42.0 44.7 19.2 11.6 5.80 6.04 6.60 6.94 8.39 7.18 6.46 6.76 8.82 9.26 6.08 3.88 1.87 1.97 1.636 1.72 1.54 2.33 1.24 1.66 1.43 1.57 0.283 0.165 5.95 6.31 6.718 7.76 9.33 7.62 6.85 6.88 9.50 10.2 7.21 4.22 0.898 0.945 1.056 1.10 1.35 1.07 0.968 1.10 1.30 1.41 1.43 0.795 4.99 5.22 6.129 6.26 7.78 5.87 5.24 6.67 7.03 7.73 8.98 4.56 0.992 1.02 1.213 1.22 1.51 1.18 1.04 1.28 1.35 1.44 1.83 0.814 2.68 2.79 3.157 3.31 4.24 3.21 2.95 3.45 3.63 4.05 5.08 2.14 0.379 0.406 0.463 0.480 0.598 0.433 0.420 0.508 0.517 0.552 0.779 0.333 2.38 2.47 2.828 2.84 3.70 2.74 2.57 3.01 3.16 3.42 4.55 2.06 0.358 0.369 0.420 0.415 0.516 0.406 0.369 0.465 0.435 0.481 0.662 0.279 7.25 7.99 6.561 7.06 8.87 9.80 6.78 6.07 2.71 3.01 2.93 2.09 0.663 0.740 0.814 1.25 1.17 0.836 1.02 1.86 0.852 1.38 1.39 1.52 5.6 6.1 9.289 11.2 13.1 8.12 12.5 13.7 25.3 23.1 31.4 26.0 3.9 4.2 5.833 5.97 7.84 4.42 7.20 5.87 12.2 12.2 6.61 4.36 0.969 1.02 1.534 1.57 2.03 1.13 1.91 1.42 2.81 2.73 5.02 5.65 0.714161 0.714587 0.713227 0.720139 0.720129 0.714831 136.41 142.45 123.15 221.25 221.12 145.92 0.511715 0.511812 0.511490 0.511671 -16.49 -14.61 -20.89 -17.37 16.706 16.813 16.484 16.778 15.002 15.017 15.005 15.023 37.051 37.091 37.238 37.171 lith. code ggu no. deg. w deg. n altitude, m sio2 tio2 al2o3 fe2o3 feo mno mgo cao na2o k2o p2o5 volatiles sum feo* mg number toc wt% s wt% zn cu ni sc v cr ga rb sr y zr nb cs ba la ce pr nd sm eu gd tb dy ho er tm yb lu hf ta pb th u isotope ratios calculated at 60 ma 87sr/86sr60 εsr 143nd/144nd60 εnd 206pb/204pb60 207pb/204pb60 208pb/204pb60 dacitelithology table 6c. chemical analyses of rocks of the nordfjord member garnet rhyolite sanidine rhyolite dacite high-ni dacite high-al graphite andesite high-p2o5 dacite high-p2o5 167 156678 alkali basalt with scarce microphenocrystic olivine and plagioclase in a very fine-grained, flow-laminated groundmass with tiny interstitial biotite. lowest flow in the nordfjord member, illuluarsuit qaqqaa, mouth of nordfjord, west disko. 327011 well crystallised basalt with scattered microphenocrysts of olivine (pseudomorphosed) and clinopyroxene and scarce sieve-textured plagioclase xenocrysts. lowest flow in the nordfjord member, top flow in the skarvefjeld profile, south disko. 332882 basalt with plagioclase glomerocrysts and microphenocrysts of olivine (mostly pseudomorphosed and rarely enclosing chromite) and scarce clinopyroxene. up to 3 mm zeolite-filled vesicles. lava flow, nunavik profile, central nuussuaq. 279298 silicic basalt with scattered microphenocrysts of olivine (pseudomorphosed) in a very fine-grained groundmass. lava flow, point 1137 m south of nordfjord, west disko. 326647 very fine-grained aphyric basalt with up to 3 mm smectite-filled vesicles. lava flow, niaqussat profile, north-west disko. 362398 strongly porphyritic basalt with glomerocrysts and phenocrysts of plagioclase, phenocrysts of olivine (pseudomophosed) and glomero crysts of calcic clinopyroxene. lava flow, qinngusaq profile, kvandalen, east disko. 279465 basalt with olivine phenocrysts (pseudomorphosed) and scarce plagioclase-clinopyroxene glomerocrysts in a fine-grained groundmass. lava flow, point 1070 m profile, outer nordfjord, west disko. 328463 porphyritic basalt with phenocrysts and up to 8 mm glomerocrysts of plagioclase, olivine (pseudomorphosed) and clinopyroxene in a very fine-grained groundmass. lava flow, uiffaq, south-west disko. 318768 porphyritic basalt with phenocrysts and up to 10 mm glomerocrysts of plagioclase, minor olivine (pseudomorphosed) and scarce clinopyroxene in a very fine-grained groundmass. lava flow, saqqaq profile, south nuussuaq. 279297 basaltic andesite, aphyric. lowest flow in the nordfjord member, point 1137 m south of nordfjord, west disko. 263934 basaltic andesite with phenocrysts of olivine (partly pseudomorphosed), scattered microphenocrysts of orthopyroxene and clinopyroxene, and scattered plagioclase xenocrysts in a glassy groundmass. lava flow, point 1114 m, vesterdalen, west disko. 318809 iron-bearing basaltic andesite with orthopyroxene-plagioclase glomerocrysts, phenocrysts of orthopyroxene, and microphenocrysts of orthopyroxene, clinopyroxene and plagioclase. up to 10 mm magma-modified sediment xenoliths with plagioclase and graphite, and groundmass with traces of native iron. centre of composite lava flow, aqajaruata qaqqaa profile, east disko. 264087 basaltic andesite transitional to andesite with scattered phenocrysts of orthopyroxene and scarce microphenocrysts of plagioclase in an extremely fine-grained groundmass. lava flow, point 1012 m profile, vesterdalen, west disko. 176626 basaltic andesite, aphyric. lava flow, enok havn, outer mellemfjord, west disko. 274430 basaltic andesite transitional to andesite with phenocrysts and microphenocrysts of orthopyroxene and plagioclase, glomerocrysts of orthopyroxene and plagioclase, and small xenocrystic aggregates of plagioclase, mullite, red spinel and graphite; traces of native iron. lava flow, ‘point 440 m northern gully’ profile, north of hammer dal, north-west disko. 176565 basalt with common olivine microphenocrysts (pseudomorphosed) and scarce plagioclase phenocrysts in a fine-grained groundmass. basal part of the mellemfjord composite lava flow, saqqarliit ilorliit, mellemfjord, west disko. 176556 native-iron-bearing basaltic andesite with phenocrysts of orthopyroxene and plagioclase and microphenocrysts of clinopyroxene, scattered xenocrystic aggregates of plagioclase, red spinel and graphite, and rare xenocryst rosettes of mullite. the groundmass is very fine-grained and contains up to 2 mm native iron-troilite bodies. lower part of the mellemfjord composite lava flow, saqqarliit ilorliit, mellemfjord, west disko. 176564 native-iron-bearing andesite with phenocrysts of resorbed plagioclase and lamellar-twinned low-ca clinopyroxene mantled by ortho pyroxene which also occurs as microphenocrysts. there are scattered xenocrystic aggregates of plagioclase with red spinel and graphite and up to 1 mm large native iron and sulphide grains in a very fine-grained groundmass. central part of the mellemfjord composite lava flow, saqqarliit ilorliit, mellemfjord, west disko. 176579 basaltic andesite with scattered small plagioclase phenocrysts in a groundmass with inhomogeneous grain size distribution and with up to 0.3 mm large ilmenite grains. lower part of composite lava flow, eqaluit (nordre laksebugt), south-west disko. 176582 native-iron-bearing basaltic andesite with phenocrysts of plagioclase and orthopyroxene (which may be mantled by lamellar-twinned pigeonite) in a well-crystallised groundmass with rusty spots after weathered native iron. there are scarce plagioclase xenocrysts with red spinel. upper part of composite lava flow, eqaluit (nordre laksebugt), south-west disko. 176448 andesite with native iron in a partly glassy groundmass. microphenocrysts of plagioclase, orthopyroxene, pigeonite and olivine, and scattered xenocrystic aggregates of plagioclase and mullite. lava flow, point 600 m profile, north side of hammer dal, north-west disko. 176473 native-iron-bearing andesite with phenocrysts of plagioclase and orthopyroxene and microphenocrysts of orthopyroxene, pigeonite and plagioclase; scattered plagioclase xenocrysts derived from sediments. lava flow c. 5.5 km east of jamma, north-west disko. 113463 native-iron-bearing andesite (transitional to basaltic andesite) with abundant phenocrysts of plagioclase and orthopyroxene and micro phenocrysts of plagioclase, orthopyroxene and pigeonite. abundant xenocrysts and xenocryst-aggregates of plagioclase, red spinel, mullite, graphite and quartz in a very fine-grained groundmass with native iron, troilite and graphite. lava flow 60 m thick, fed from crater, niaqussat profile, north-west disko. 176411 native-iron-bearing andesite with phenocrysts of plagioclase and orthopyroxene and a large range of microphenocrysts and xenocrysts (pedersen 1981 table 2 no. 1 shows a detailed modal analysis). there are decimetre-sized cognate inclusions of norite. lava flow within the niaqussat crater, niaqussat profile, north-west disko. 264064 native-iron-bearing andesite with common phenocrysts of plagioclase and orthopyroxene and microphenocrysts of plagioclase, ortho pyroxene and clinopyroxene and possibly olivine pseudomorphs. the groundmass contains up to 0.5 mm grains of native iron. lava flow, point 1300 m profile, vesterdalen, west disko. 264067 native-iron-bearing andesite with an inhomogeneous groundmass with variable grain size. scarce up to 1 mm phenocrysts of plagioclase, orthopyroxene and clinopyroxene and common prismatic microphenocrysts of orthopyroxene and clinopyroxene. scarce xenocrysts include quartz, plagioclase and mullite. lava flow, point 1300 m profile, vesterdalen, west disko. 176441. native-iron-bearing dacite (transitional to andesite) with common phenocrysts of plagioclase and orthopyroxene and microphenocrysts of plagioclase, orthopyroxene and clinopyroxene. abundant aggregates of plagioclase and quartz and aggregates of plagioclase, spinel and graphite. xenocrysts include plagioclase, quartz, spinel, cordierite, corundum and mullite. base of 120 m thick lava flow, point 600 m profile, north side of hammer dal, north-west disko. phenocryst phases as observed in thin section are mentioned in order of decreasing abundance. samples of subaerial lava flows are usually taken in massive columns a few metres above the flow base. table 6d. notes on analysed samples of the nordfjord member 168168 176466 native-iron-bearing dacite with common phenocrysts of plagioclase and orthopyroxene and scarce ilmenite with armalcolite and rutile formed by progressive reduction and sulphidation. the very fine-grained groundmass contains native iron and troilite. a detailed mode, which also shows the range of xenocrysts, is given by pedersen (1981 table 2 no. 3). lava flow, jamma, coast of north-west disko. 176471 the most silicic native-iron-bearing volcanic rock in the nuussuaq basin. dacite with common phenocrysts of plagioclase, orthopyroxene and pigeonite; microphenocrysts of ilmenite have been transformed to aggregates of ilmenite, rutile, native iron and troilite. the very fine-grained groundmass contains native iron, troilite and graphite. the abundance of various xenocrysts and cognate microxenoliths with plagioclase, orthopyroxene and pigeonite is described in pedersen (1981). lava flow, jamma, coast of north-west disko. 176555 native-iron-bearing dacite with scattered small phenocrysts of plagioclase and orthopyroxene and abundant microphenocrysts of tridymite and plagioclase in a fine-grained groundmass with scarce native iron. there are scattered xenocrystic aggregates of quartz and plagioclase and rare spinel. lava flow, saqqarliit ilorliit, mellemfjord, west disko. 176443 native-iron-bearing dacite with abundant phenocrysts of plagioclase and orthopyroxene and microphenocrysts of ilmenite and armalcolite. abundant xenocrysts and aggregates of plagioclase, quartz, cordierite, spinel and graphite; common xenoliths of picrite with chromite, and common cognate fine-grained norites. the very fine-grained groundmass contains native iron, troilite and graphite. lava flow fed from the point 440 crater, point 600 m profile, north side of hammer dal, north-west disko. 326550 high-al graphite dacite with disseminated graphite and phenocrysts of orthopyroxene and plagioclase in a fine-grained groundmass. abundant xenocrysts of mullite, cordierite, aluminous spinel and corundum. rounded cobble from conglomerate in nordfjord member from the ‘point 440 m northern gully’ profile north of hammer dal, north-west disko. 326465 garnet rhyolite with graphite flakes in the groundmass. more than 20% phenocrysts of plagioclase, quartz, orthopyroxene and garnet and minor ilmenite, apatite and zircon. abundant xenocrysts include plagioclase, quartz, sillimanite, corundum and hercynite. large (60 cm × 40 cm) rounded lava boulder in conglomerate. sedimentkløften, south side of hammer dal, north-west disko. 156518.1 garnet rhyolite with graphite flakes in the groundmass. about 25% phenocrysts of plagioclase, quartz, orthopyroxene, garnet and biotite and minor ilmenite, apatite and zircon. there is a range of sediment xenoliths, xenocrysts and cognate clusters of plagioclase, orthopyroxene, pigeonite and almandine-rich garnet. rounded lava cobble in conglomerate. sedimentkløften, south side of hammer dal, north-west disko. 156516 sanidine rhyolite, pitchstone with tiny flakes of graphite in the groundmass glass. about 5% phenocrysts of plagioclase, quartz, sanidine, biotite and very minor ilmenite, zircon, apatite, monazite and almandine-rich garnet. there are very scarce aluminous xenocrysts and highly equilibrated sediment xenoliths composed of aggregates of feldspar, hercynite, orthopyroxene, ilmenite, sillimanite and graphite. rounded lava cobble in conglomerate. sedimentkløften, south side of hammer dal, north-west disko. 156559 sanidine rhyolite, vesiculated glassy rock with <10% phenocrysts of plagioclase, quartz, sanidine and biotite and very minor ilmenite, graphite and garnet. lithic clast in pumice tuff. north side of outer rink dal, north-west disko. table 6e. notes on analysed samples of the nordfjord member rock type ggu no. toc s fe0 sio2 andesite flow at niaqussat 176411* 0.09 0.41 2.4 60.71 dacite clast n of hammer dal 176486 0.9 0.22 0.9 65.49 dacite clast n of hammer dal 326550 2.72 0.02 n.a. 63.52 dacite clast n of hammer dal 326551 2.60 0.02 n.a. 62.75 hammer dal point 600 m, dacite lava flow 176442 n.a. 0.2 1.6 63.07 point 600 m, dacite lava flow 176443* 1.17 0.64 + 66.03 point 440 crater, dacite 176497 0.8 0.67 + 66.09 jamma dacite lava 176466* 0.07 0.22 0.7 64.64 dacite lava 176471* 0.27 0.28 0.14 67.71 mellemfjord composite lava flow lower part, basaltic andesite 176556* n.a. 0.22 + 54.72 central part, andesite 176564* n.a. n.a 0.26 60.67 central part, andesite 176548 n.a. 0.26 + 61.19 sedimentkløften, hammer dal sanidine rhyolite clast in conglomerate 113515 0.20 n.a. 0 75.68 garnet rhyolite clast in conglomerate 113508 0.53 n.a. 0 70.76 garnet rhyolite clast in conglomerate 156518.1* 0.17 n.a. 0 72.68 sediment xenolith in rhyolite clast 156518.2** 11.72 n.a. 0 52.47 airfall tuffs north of hammer dal white rhyolitic pumice clast in tuff 326497† 0.51 0.01 0 n.a. dark grey rhyolitic pumice clast in tuff 326497† 3.1 0.02 0 n.a. toc: total organic carbon. n.a. : not analysed. +: trace amounts. *sample included in table 6 (major and trace elements, some isotopes). **sample included in table 13 (major and trace elements, some isotopes). †: hand-picked pumice clasts from tuff with a mixed population of white and dark grey pumice clasts. total sulfur determined by combustion using a leco cs-200 induction furnace apparatus.total organic carbon determined similarly after elimination of carbonate-bonded carbon through several stages of prolonged treatment with hot hydrochloric acid (hcl, 2n). analyst: j. bojesen-koefoed. metallic iron determined by the hgcl2•nh4cl method. analyst m. mouritzen. sio2 is recalculated on volatile-free basis except for the carbon-rich sediment xenolith. table 7. chemical analyses of carbon, sulfur and metallic iron in rocks of the nordfjord member 169 niaqussat member summary of the main features of the niaqussat member • original extent over all of disko and southern and eastern nuussuaq, with depocentre on western disko. • except for a thin soil at some localities, the member directly overlies the nordfjord member; the lava flows covered the cratered landscape on north-west disko and formed a new level surface. • the onset of the member indicates renewed upflow of magnesian magmas from depth. • the lower niaqussat member is dominated by picrites and magnesian basalts. there are a few magnesian basaltic andesites, some of which are native-iron-bearing. the middle and upper niaqussat member comprise successively more evolved silicic basalts. • all igneous rocks are more or less crustally contaminated; sediment xenoliths are scarce but in one case occur in abundance. • dykes with compositions similar to weakly contaminated niaqussat member basalt are found over large areas of disko and nuussuaq, suggesting eruption over wide areas. • on north-east nuussuaq, basalt flows of the upper niaqussat member extended the volcanic plateau c. 10 km eastward over the elevated gneiss country and invaded sediments with tuff layers believed to correlate with the nordfjord member. • the two highest preserved lava flows of the upper niaqussat member on north-east nuussuaq have a hightio2, slightly enriched chemical composition very similar to that of the two highest flows on west disko, suggesting that the member on nuussuaq is close to completely represented. lithostratigraphy of the niaqussat member revised member history. the niaqussat member was informally established by pedersen (1975a) on north-western disko. the definition is formalised here and extended to cover the whole of disko and nuussuaq. name. after niaqussat, a coastal slope on north-western disko (fig. 4). distribution. the niaqussat member originally extended over the whole of disko and eastern nuussuaq; its main present occurrence is in the downthrown, west-dipping fault blocks on western disko. it has been removed by erosion over wide areas on central and southern disko but is preserved on peaks and ridges on eastern and north-eastern disko (e.g. pedersen et al. 2001; south disko section; central disko section). it has been eroded away on southern and central nuussuaq but is present at high altitudes on north-eastern nuussuaq (larsen & pedersen 1992; central nuussuaq section). type section. the type section is composite. for the lower and middle niaqussat member: point 440 m, northern gully, on north-western disko (fig. 14, profile 11; figs 103, 104a). the locality is an e–w-oriented gully c. 4 km north of the river in hammer dal. the niaqussat member overlies the nordfjord member and is exposed over a lateral distance of c. 500 m at around 300 m altitude; the succession is faulted and tilted and dips 22°w (photogrammetrically measured section in fig. 104a). the upper niaqussat member is eroded away at this locality; for this part, the type section is the sapernuvik profile above c. 140 m, north of outer kangerluk, western disko (fig. 14, profile 1; see also figs 161, 171). reference sections. ikorfarsuit, outer mellemfjord, western disko, above c. 400 m (fig. 14, profile 3). the high peak pyramiden, at above 1830 m, northern disko (fig. 10, profile 3). the southern shoulder of the mountain qinngusaq above the innermost part of kvandalen, above c. 1290 m, eastern disko (fig. 10, profile 8; central disko section at 81–83 km). the north-western wall of frederik lange dal, at above c. 1150 m, eastern disko (fig. 10, profile 9). the southern slopes (nunavik), at above c. 1680 m, of the mountain point 2000 m, eastern nuussuaq (fig. 12, profile 13; central nuussuaq section at 67–68 km). thickness. on western disko, the niaqussat member attains thicknesses of up to 500 m at sapernuvik where it is most complete; more commonly it is 200–300 m thick, dependent on the level of erosion. thicknesses are up to 160 m on north-western disko, up to 180 m on eastern disko and up to 200 m on eastern nuussuaq. the maximum number of flows preserved in one section is c. 30 at 170170 sapernuvik, c. 10 at frederik lange dal on eastern disko, and at least seven on eastern nuussuaq. the youngest flows preserved occur on eastern nuussuaq and at sapernuvik on western disko. lithology. the niaqussat member consists of subaerial lava flows. these comprise olivine-phyric picrites and olivine-plagioclase-phyric and aphyric magnesian basalts and normal basalts. many of the more magnesian lavas show a characteristic flow folding and flow lamination (fig. 150). a few flows are basaltic andesites, some of which carry native iron. subdivisions. the niaqussat member has been subdivided into three informal units (units 530, 531 and 532), for convenience also called the lower, middle and upper niaqussat member. the subdivision is based on a stepwise decrease in mgo contents up-section; the units are clearly separated in geochemistry diagrams (see also figs 167, 168), but the middle unit is generally indistinguishable in the field. boundaries. the niaqussat member conformably overlies the nordfjord member. its lower boundary is placed at the base of a succession of lithologically distinctive, flow-folded pahoehoe lavas of olivine-phyric magnesian basalt and picrite. in places, there is a small sediment horizon at the base, but in most cases the niaqussat member lava flows rest directly on lavas and crater deposits of the nordfjord member. particularly on north-western disko, the top of the nordfjord member formed a very irregular and intensely cratered volcanic landscape which was drowned by the highly fluid lava flows of the niaqussat member so that an even lava field was re-established. this is for example seen at the point 440 crater site (fig. 143). elsewhere, the flows continued to build up the existing lava field. the upper boundary is erosional except in a small area at sapernuvik on westernmost disko, where the niaqussat member is conformably overlain by the lava flows of the sapernuvik member. age. paleocene, 61–60 ma, magnetochron c26r, based on radiometric dating (storey et al. 1998; larsen et al. 2016). correlation. none certain. lower niaqussat member (unit 530) composition and petrography. the onset of the niaqussat member volcanism is marked by a return to olivinephyric, relatively primitive, picritic magma compositions. unit 530 comprises picrites with up to 15 wt% mgo and olivine-microphyric magnesian basalts with more than 7 wt% mgo (except for three samples with 6–7 wt% mgo). there are also a few more silicic flows of basaltic andesite, some of which carry native iron. distribution and thickness. unit 530 constitutes the major volume of the niaqussat member. the thickest and largest occurrences are on western disko between mellemfjord and hammer dal (fig. 14), where thicknesses are close to 200 m and up to 240 m. erosional remnants are widespread on peaks and ridges in other parts of disko and on north-eastern nuussuaq with thicknesses up to 100 m. on nuussuaq unit 530 has only been sampled in one profile at nunavik where a c. 30 m thick succession remains (fig. 12); from photogrammetric studies (central nuussuaq section), the unit appears to comprise only a few flows in this area. mf142_8cmakp_1972_has_17_06cshmod_c.eps 0.5 m fig. 150. picrite lava flow of the lower niaqussat member showing the characteristic flow lamination and folding of this part of the member. note compass in lower right corner. 171 dykes with chemical compositions similar to the lower niaqussat member are known from both sides of the vaigat strait on eastern disko and south-eastern nuussuaq, as well as on the south coast of disko (larsen & pedersen 1992; see also fig. 174 and table 8). this suggests that the eruption sites were widespread. lithologies. the magnesian basalt to picrite flows occur as successions of pahoehoe lava flows from a few metres to more than 30 m thick. many of the flows display a characteristic flow folding (fig. 150) due to an inhomogeneous distribution of vesicles, which has given rise to grain-size variation throughout the flows. despite a universal chemical signature of crustal contamination (see below), the typical picrites and olivinemicrophyric basalts of unit 530 are with few exceptions devoid of sediment xenoliths or xenocrysts. the few strongly contaminated lava flows, some of which carry native iron, contain numerous sediment xenoliths. western disko on western disko, unit 530 forms a thick succession of many <10 m thick pahoehoe lava flows (figs 151, 152), which seem to have built up small shield volcanoes. feeder dyke and lava flows on the northern wall of rink dal. an up to 4 m thick dyke of olivine-microphyric basalt is situated in the north-western corner of rink dal about 1.5 km east of the coast. the dyke strikes e–w over a distance of a few hundred metres. it is associated with a local up-doming of similar basalt and of a layer up to 5 m thick of partly welded tuff of olivine-microphyric basalt. the deposit is interpreted as having erupted from lava fountains along an eruptive fissure fed from the dyke, emitting pahoehoe lava flows and tuffs. the northern wall of rink dal is strongly faulted and not well exposed (pedersen 1975a, plates 1, 2), but an excellent section through unit 530 is exposed about 3–4 km east of the feeder dyke (fig. 103, loc. 16; fig. 151). here several strongly plagioclase-phyric basalts of the mf143_17cma1985_07_04mod_c.eps 530 basalt 520 530 picrite 50 m 530–532 d 530 fe-ba fig. 151. well-exposed lava succession of the niaqussat member. the lower niaqussat member (unit 530) comprises, lowest, a basaltic tuff layer (1.5 m thick, not seen in the picture), an olivine-microphyric basalt flow 15 m thick, two native-iron-bearing basaltic andesite flows (fe-ba) and a characteristic succession of many thin picrite pahoehoe flows, c. 30 m thick. the thin picrites are overlain by thicker basalt flows of units 530–532. unit 530 is underlain by tuffs and basalt flows of the nordfjord member (520). the succession is cut by a younger dyke (d). north side of the western rink dal, western disko. for locality, see fig. 103, loc. 16. 172172 nordfjord member are covered by a tuff partly altered to laterite, which defines the top of the member. the laterite is covered by an up to 1.5 m thick olivine-microphyric basaltic tuff assigned to the niaqussat member. the tuff has bluish black and reddish, fiamme-like tongues and is similar in structure to the basaltic tuff overlying the feeder dyke described above. the tuff is covered by a c. 15 m thick, flow-folded and flow-laminated olivine-microphyric basalt with 10.4 wt% mgo; this is overlain by a >6 m thick basaltic andesite flow with 9.7 wt% mgo, with mf144_11cmakp_25_c1_79mod_c.eps fig. 152. a characteristic, thick succession of many thin picrite pahoehoe flows of unit 530 at qasigissat, west disko (fig. 14, profile 6). height of section in view 170 m. fig. 153. accumulation of xenoliths at the base of a lava flow of unit 530. the closepacked xenoliths are rounded, up to 20 cm in size and consist of strongly magma-equilibrated mudstone that is now transformed into plagioclase-spinel-orthopyroxene rocks. the red arrows point to two xenoliths such as shown in fig. 154b. length of hammer 47 cm. illuluarsuit qaqqaa, westernmost rink dal. for locality, see fig. 103, loc. 18. the occurrence is exposed in a small gully on the southern side of a fault block with lavas dipping 10°w. mf146_11cmakp_1972_h_04_17igdlkxenmod2_c.eps 530 lava 173 native iron and sulfides in the basal part and abundant mudstone xenoliths up to 15 cm in size concentrated in the upper part. this flow is covered by a basaltic andesite flow several metres thick, with rusty spots and sediment xenoliths. then follows a spectacular succession of c. 15 contaminated, picritic pahoehoe lava flows each a few metres thick (fig. 151), on top of which there are several >5 m thick flows of olivineand slightly plagioclase-phyric basalts. the monotonous succession of thin picrite flows suggests proximity to a feeder system in the western part of rink dal, which produced small lava shields. similar distinct successions of very thin, olivine-rich pahoehoe flows of unit 530 also occur along the west coast of disko for 25 km farther south from rink dal to qasigissat (fig. 14, profiles 6, 7; fig. 152), whereas fewer and thicker olivine-microphyric basalt flows are encountered farther south and east of the coast (fig. 14, profiles 3–5). xenolith accumulation in olivine-microphyric basalt lava in rink dal. an exception to the general lack of sediment xenoliths is found in western rink dal (fig. 103, loc. 18), where a thick accumulation of xenoliths occurs at the base of a lava flow (fig. 153). at this locality, a strongly plagioclase-phyric basalt flow covered by a laterised, 30 cm thick tuff marks the top of the nordfjord member. this is overlain by an olivine-microphyric basalt tuff c. 0.8 m thick, with fiamme structure very similar to the basal niaqussat member tuff 4 km to the north-west described above. the tuff is covered by an olivine-microphyric basalt flow 9–17 m thick (fig. 154a) with very prominent flow lamination and flow folding structures (fig. 150). the lowermost part of this flow is an up to 2.5 m thick deposit of rounded, up to 20 cm large xenoliths of strongly magma-equilibrated mudstones transformed into plagioclase-spinel-orthopymf145a_8cm156677dglmodcut_c.eps ol ol 5 mm a pl mf145b_8cm156676xdglmodcut_c.eps 5 mm b pl ca pl opx opx opx sp sp fig. 154. thin sections (scanned) of an olivine microphyric basalt flow packed with mudstone xenoliths. a: flow-laminated basalt with microphenocrysts of olivine (ol) and plagioclase (pl) in a groundmass with zones of variable grain size. b: pyrometamorphosed mudstone xenolith now consisting of an aggregate of plagioclase (pl, predominant), spinel (sp) and orthopyroxene (opx). vesicles are filled with calcite (ca). see text for details. sample 156677, rink dal, western disko. 174174 roxene rocks rich in plagioclase (fig. 154b). the xenoliths are packed in a matrix of vesiculated basalt, and the deposit can be followed for 300 m along the base of the flow. on the northern part of the same fault block the lava flow is exposed again c. 1 km to the north of the xenolith-rich site, but here without visible xenoliths. the occurrence links the lavas of unit 530 to high-level magma reservoirs within sediments, and the huge accumulation of sediment xenoliths must indicate close proximity to another erup tion site in the lowermost part of the niaqussat member. eastern disko olivine-microphyric basaltic to picritic pahoehoe lava flows of unit 530 cap many plateaus and mountaintops on eastern disko (pedersen & larsen 1987). this is well exemplified by point 1123 m on the southern side of kvandalen (fig. 7, profile 12; fig. 155; also larsen & pedersen 1988, fig. 4). here, two basaltic lava flows of the nordfjord member are covered by a lateritic soil layer a few decimetres thick and overlain by three olivine-microphyric magnesian basalt lavas of unit 530. some of the lava flows of the niaqussat member were erupted locally, as attested by the occurrence of basaltic dykes of similar composition in the area (larsen & pedersen 1992, table 1, no. 8, sample 318824 with 12.64 wt% mgo from akunneq). unit 530 is 90 m thick at frederik lange dal (fig. 10, profile 9) and c. 150 m thick at point 1530 m north-west of kvandalen (fig. 10, profile 6). the lower part of unit 530 is dominated by olivine-microphyric to almost aphyric magnesian basalts, which are macroscopically almost devoid of sediment xenoliths or xenocrysts and carry neither native iron nor sulfides. however, at a level 40–55 m above the base of the member in frederik lange dal and 25–30 m above the base at qinngusaq mountain, lava flows, which are clearly contaminated, occur in both profiles. this part of the profile at point 1530 m (fig. 10, profile 6) remains unsampled. compound lava flow with sediment xenoliths and sulfides in frederik lange dal. about 40 m above the base of the niaqussat member in frederik lange dal (fig. 10, profile 9), a compound lava flow consisting of at least four flow lobes overlies a thin horizon of lateritic soil which in turn overlies three lava flows of olivine-microphyric mf147_17cmakp_1987_11_12mod2_c.eps 530 520 517 fig. 155. the top of point 1123 m on the south side of kvandalen, eastern disko (fig. 7, profile 12). the lowest, thick lava flow belongs to unit 517 of the rinks dal member. this is followed by two thinner flows of the nordfjord member (unit 520); these are covered by a few decimetres thick, brick-red lateritic soil layer and overlain by three olivine microphyric basalt flows of the lower niaqussat member (unit 530). the shown succession is 130 m thick. 175 basalt. the compound flow is 12–15 m thick and has a rusty brown weathering colour (fig. 156). the lowermost flow lobe is composed of olivine-microphyric basalt (51 wt% sio2, 11 wt% mgo); its lower part contains disseminated iron sulfides (pyrrhotite or troilite) and scattered, few-millimetres-sized xenoliths of magma-modified mudstone with fine-grained plagioclase and reddish spinel which form radiating aggregates derived from reaction with mullite (as in fig. 109a and in pedersen et al. 2017, fig. 111). the flow must have been erupted locally on eastern disko. basaltic andesite lava flow with native iron and sulfide at qinngusaq. a c. 4 m thick, light grey aa lava flow is exposed about 25 m above the base of the niaqussat member at qinngusaq (fig. 10, profile 8). it is a highly magnesian basaltic andesite (55.8 wt% sio2, 9.2 wt% mgo) with scattered up to 8 mm aggregates of native iron and sulfide (fig. 157). the rock contains resorbed olivine phenocrysts and abundant orthopyroxene phenocrysts and microphenocrysts. there are also scattered xenocrysts of quartz and abundant aggregates of sievetextured plagioclase derived from sediment–magma reaction. the flow also contains centimetre-sized, magma-modified mudstone xenoliths. the flow is the only native-iron-bearing lava flow in the niaqussat member on eastern disko and nuussuaq. the chemical composition of the flow is comparable to that of the basaltic andesite dyke with native iron and sulfide at illukunnguaq (pauly 1958) at the vaigat coast 5 km north-east of qinngusaq, and the dyke is therefore considered to be a likely feeder for the lava flow (larsen & pedersen 1992, dyke analysis table 1, no. 7, sample 362140). eastern nuussuaq lava flows of the niaqussat member occur in the elevated country east of the eastern boundary fault on north-eastern nuussuaq. here, between points 2080 m and 2000 m, there is a less than 50 m thick succession of 2–5 light grey picritic pahoehoe flows of unit 530 (fig. 12, profile 13 (nunavik); figs 158, 159). the flows are separated from the underlying two basalt flows of the nordfjord member by 1–2 m of yellow brown tuffaceous volcaniclastic sandstone with up to 1 cm large fragments of plant mf148_17cmakp_1987_07a_13mod_c.eps xs 530 520 518 517 531–532 fig. 156. part of the sampled profile in frederik lange dal, eastern disko (fig. 10, profile 9). unit numbers are indicated. the compound flow of unit 530 with sediment xenoliths and sulfides is labelled xs. the shown succession is 350 m thick; there is a strong perspective shortening uphill, compare with fig. 10, profile 9. for location, see fig. 140. 176176 fossils. traces of similar sandstone are also seen in the scree coming down from the top of the picritic succession. about 3–4 km east of the nunavik profile both the picrites and the underlying lava flows of the nordfjord and rinks dal members are banked up against a topographic high of precambrian gneiss that blocked further eastward progression (central nuussuaq section at 71.6 km; see also fig. 163). middle niaqussat member (unit 531) composition and petrography. unit 531 comprises a few lava flows of silicic basalt with 50–52 wt% sio2, 6–7 wt% mgo and 1.7–2.2 wt% tio2. the rocks show only little petrographical variation; most have plagioclase phenocrysts and glomerocrysts varying in size between 0.5 and 2 mm, and in addition phenocrysts and glomerocrysts of augite <0.5–1 mm in size. olivine is scarce or absent. some rocks are almost aphyric. xenocrysts of sedimentary origin have not been observed; however, occasional aggregates of plagioclase xenocrysts up to 7 mm in size with sieve-textured cores provide evidence for reaction between magma and crustal rocks. unit 531 is chemically well defined but is not recognisable as a field marker horizon. distribution and thickness. flows of unit 531 overlie the more magnesian flows of unit 530 in four profiles on western disko and three profiles on eastern disko. the unit is absent in an area in vesterdalen on western disko mf149_8cm362179polmod_c.eps ve xen opx fe+tro 5 mm fig. 157. thin section (scanned) of a magnesian basaltic andesite of unit 530. the vesicular (ve) groundmass contains orthopyroxene microphenocrysts (opx) and a large aggregate of native iron and troilite (fe+tro). the xenolith (xen) is a plagioclase-rich, strongly magma-modified mudstone. sample 362179, qinngusaq, eastern disko. gv03_03_062_lml 65 66 67 68 69 70 km 200 400 600 800 1000 1200 1600 2000 m i2 n1 mnmm mu n1 mn mm mumu gn gn e e e e e e e e e e ee e i1b point 2000 m 532 520 517 517 450 520+530 w e fig. 158. photogrammetrically measured section of the mountainside between nunavik (just left of the picture) and point 2000 m, eastern nuussuaq. blue three-digit numbers are lithological codes; other annotations as in the original. e: entablature lava. the red horizon labelled mn comprises both unit 520 and unit 530. unit 531 is not present. compare with fig. 159. excerpt from the central nuussuaq section (pedersen et al. 2002a). 177 and at pyramiden on northern disko. in many other profiles, the level is eroded away. its presence has not been confirmed on nuussuaq. the unit is 15–35 m thick and only comprises one to three flows except at point 440 m in the hammer dal area, where there are five flows with a combined thickness of 35 m, truncated upwards by erosion. at one locality (fig. 14, profile 2), 10 cm of lateritic tuff or soil separate unit 531 and 530, but otherwise there are no indications of significant interruption in volcanic activity between the two units. upper niaqussat member (unit 532) composition. unit 532 comprises lava flows of silicic basalt with 50–52 wt% sio2, 4.9–6.4 wt% mgo and 2.4– 2.9 wt% tio2, which are capped by two or more flows of enriched basalt with c. 3.9 wt% tio2 on nuussuaq (fig. 12, profile 12) and by two flows of very enriched basalt with 4.3 and 5.2 wt% tio2 at sapernuvik on south-western disko (fig. 14, profile 1). petrography. the basalts of unit 532 vary petrographically from rocks with distinctive millimetre-sized glomerophyric clusters of plagioclase, augite and pseudomorphed olivine (fig. 160a) to almost aphyric rocks. some of the more silicic basalts show flow lamination; besides scarce plagioclase phenocrysts and glomerocrysts, these basalts contain prismatic and commonly twinned microphenocrysts of clinopyroxene, some of which may be subcalcic or pigeonitic. in rare cases, these flows contain stellate clusters of prismatic clinopyroxene. the enriched basalt flow no. 2 from the top of the niaqussat member at sapernuvik on disko with 4.3 wt% tio2 has scarce glomerocrysts of plagioclase up to 4 mm in size and scattered microphenocrysts of plagioclase, augite and pseudomorphed olivine. the top flow with 5.2 mf151_17cmakp_1991_11_04udmod_c.eps 517 520 530 532 tuff fig. 159. lava flows east of nunavik, eastern nuussuaq (fig. 158 at c. 67 km). the rinks dal member is present as unit 517; the nordfjord member (unit 520) comprises two basalt flows, the lower niaqussat member (unit 530) comprises 2–5 light grey picritic pahoehoe flows and the upper niaqussat member (unit 532) comprises about seven dark brown basalt flows. the combined thickness of units 520 + 530 is about 60 m. for location, see fig. 163. 178178 wt% tio2 is a fine-grained basalt with thin, platy plagioclase phenocrysts up to 2 mm in size, scattered augites up to 0.5 mm in size and fairly abundant olivine microphenocrysts up to 0.3 mm in size (fig. 160b). many olivine microphenocrysts have an oxidised rim and an unaltered core. despite the very high tio2 content, microphenocrysts of iron-titanium oxide are not seen. distribution and thickness. unit 532 has been sampled in five profiles on western disko, one on northern disko (pyramiden), one on eastern disko (frederik lange dal), and two profiles and a reconnaissance site on eastern nuussuaq. it is most complete at sapernuvik where ten flows (160 m) are present. at frederik lange dal, five flows (c. 60 m) are present. in other profiles on disko, only one flow of this unit has been left by erosion. the unit is in fact better preserved on north-eastern nuussuaq, where there are up to seven flows with a combined thickness of around 200 m and up to 300 m in the easternmost parts of nuussuaq where flows have ponded in lows on the gneiss surface. disko upper niaqussat member at sapernuvik on western disko. the most complete exposures, and the only ones where the top of the niaqussat member is preserved, are found in the area between sapernuvik and kingittup qaqqaa on western disko. the upper niaqussat member is here c. 230 m thick (fig. 14, profile 1) and consists of about 12 robust flows with brownish weathering colour but becoming more greyish towards the top of the succession (fig. 161). thin horizons of lateritic soil occur between most flows, indicating a low eruption frequency for this part of the volcanic succession (fig. 162; fig. 14, profile 1). the area has been briefly described by pedersen (1977b); it was mapped by pedersen (1977b, plate 1) mf152a_8cm332900polmod_c.eps ol 5 mm cpx pl a mf152b_8cm176615polcutmod_c.eps ol ol ol 5 mm pl b fig. 160. thin sections (scanned) of basalts of the upper niaqussat member. a: typical basalt with millimetre-sized, glomerophyric clusters of plagioclase (pl), pseudomorphed olivine (ol) and sparse augite (cpx) in a fine-grained groundmass. sample 332900, point 2080 m, nunavik, eastern nuussuaq. dated sample (storey et al. 1998). b: the ti-rich uppermost flow of the niaqussat member: a fine-grained basalt with thin platy plagioclase (pl) phenocrysts, olivine microphenocrysts (ol) and scattered augite. sample 176615, sapernuvik, west disko. 179 and pedersen & ulff-møller (1987) and is shown on the south disko section at 0–5 km (see fig. 171). the top of the niaqussat member, which ends with two unusually ti-rich flows, is best exposed on the steep corrie wall c. 3 km south-west of kingittup qaqqaa (fig. 162). the 20–30 m thick, second highest flow with 4.3 wt% tio2 here overlies a ‘normal’ flow with 2.54 wt% tio2, separated by a decimetre-thick lateritic soil layer. the flow can be observed as a local marker horizon in several corrie walls between sapernuvik and kingittup qaqqaa (fig. 161); it must once have covered an area of more than 16 km2 and represents an original volume of at least 0.5 km3. the flow is capped by a thin lateritic soil layer, which is covered by two lava lobes (considered to be one flow) with a maximum thickness of c. 13 m and containing 5.2 wt% tio2. the volume of this flow cannot be assessed accurately but must be well below 0.1 km3; it must have been erupted close to the vicinity of sapernuvik. a less than 20 cm thick horizon of lateritic soil on top of the flow separates it from the olivine-phyric basalts of the sapernuvik member. there is no indication of prolonged volcanic quiescence between the formation of the niaqussat and sapernuvik members, despite a very marked contrast in their geochemical signatures. eastern nuussuaq lava flows of unit 532 form a succession more than 200 m thick in the high area north and north-east of nunavik (larsen & pedersen 1992, figs 3, 4). the lavas rest on basalt flows of the lower niaqussat member and, in the eastern parts, on precambrian gneiss. the flows of unit 532 are the youngest preserved volcanic rocks on nuussuaq east of the itilli fault; their upper boundary is erosional. about seven lava flows are preserved above the picrites of unit 530 (fig. 12, profiles 12, 13). these basalts (the upper lava sequence of larsen & pedersen 1992) are very similar to the flows of unit 532 elsewhere. one of the flows (sample 332900) was dated by storey et al. (1998) fig. 161. the lava succession in the type section for the upper niaqussat member and the sapernuvik member. the sample profile runs along the ridge in the right foreground. the niaqussat member is well developed and comprises all of the lower (530), middle (531) and upper (532) parts of the member. it is overlain by the three preserved flows of the sapernuvik member (540). the top plateau is at around 600 m altitude. photogrammetric interpretation in fig. 171. paakkarut, sapernuvik, west disko. mf153_17cml9119_1_06mod2cut_c.eps 540 kingittup qaqqaa 531 520 530 532 d 540 180180 to 60.2 ± 0.5 ma (recalculated), i.e. selandian. the lava flows have a characteristic dark brown to almost black weathering colour (fig. 159). they range in thickness from less than 5 m to more than 50 m; several have very prominent colonnades and entablature zones, indicating emplacement in a wet environment. a 2 m thick horizon of laterised, red volcaniclastic sediment is exposed in the upper part of the succession, indicating a distinct time gap in the volcanism; several more soil horizons may exist hidden in scree. east of a palaeo-topographic high of precambrian gneiss near point 2000 m (central nuussuaq section at 71.6 km), the lava flows of unit 532 flowed directly onto the gneiss (fig. 163). in this area the flows ponded in small basins where they invaded a thin succession of paleocene quartzo-feldspathic clastic sediments (cenmf154_11cma1974_sapernuvik4no10uglasm2_c.eps 532 530–531 lat lat lat 532 (176616) 540 176615 fig. 162. the lava succession in a steep corrie wall between sapernuvik and kingittup qaqqaa, west disko. the upper niaqussat member (unit 532) here comprises nine lava flows with lateritic soil (lat) horizons between them. the 20–30 m thick, second highest flow (with sample 176616) is a local marker horizon. the highest niaqussat member flow is compositionally unusual, with the highest tio2 content of all lavas in the maligât formation (sample 176615: 5.1 wt% tio2). it is overlain by the sapernuvik member flows (540). the two numbered samples were taken in the respective flows although not at this locality; the actual sample profile is shown in fig. 161. 52°52°30' 70°29' 5 km gv01_02_161_lml.eps 5 km 1985 0.1 2.0 0.6 0.40.4 pc pc pcnunavik nuugajukassak alleq nuugajukassak qulleq makittarissagaq um iar to rfi up s er m ia pc mni1 vo?vo vo vo mni3 mni3 t3c mf2 mf2 2011 1715 fig. 159 fig. 163. map demonstrating the eastward onlap of successively younger volcanic units onto the high gneiss terrain east of nunavik on eastern nuussuaq. the gneiss (pc) is pale red; grass green (vo) is the vaigat formation, dark green (mf2) is the upper maligât formation unit 517, olive green (mni1) is the combined nordfjord and lower niaqussat members, and dark purple (mni3) is the upper niaqussat member (unit 532). the red arrow points to the location of paleocene sediments (t3c) and figs 164–166. excerpt from the 1:100 000 scale special map sheet paatuut (pedersen et al. 2007a). annotations from the original. 181 tral nuussuaq section at 59–79 km; fig. 164; pedersen et al. 2007a; fig. 12, profiles 12, 14). these small basins may have reflected the local topography and may have been short-lived. the sediments may be referred to the atanikerluk formation, but it is uncertain whether they represent the youngest and most easterly remnants of the assoq member (dam et al. 2009; pedersen et al. 2007a, unit t3c). the lava flows can be followed for about 10 km towards the east but not continuously, and most contacts are obscured by scree or ice. at makittarissagaq the succession is at least 200 m thick and consists of at least five flows, with traces of sediment between the two lowest flows (central nuussuaq section at 78.3 km). the easternmost localities with niaqussat member lava flows are fig. 164. photogrammetrically measured section of some of the easternmost exposures of lava flows on eastern nuussuaq, 3 km north of makittarissagaq (fig. 163). the flows (n1) belong to the upper niaqussat member and have invaded tuff-bearing sediments (t2) assigned to the atanikerluk formation. excerpt from the central nuussuaq section, panel l–m (pedersen et al. 2002a). annotations from the original. 1200 1400 1600 500 1000 m n1 n1 n1 t2 t2 gn in in gv03_03_063_lml w e mf157_17cmolieagat_4_04mod2_c.eps 532 inv 532 inv 532 inv 100 m fig. 166 sed fig. 165. lava flows of the upper niaqussat member (unit 532) invading tuff-bearing sediments (sed) assigned to the atanikerluk formation. see photogrammetric interpretation in fig. 164 and location in fig. 163. the sediments are shown in more detail in fig. 166. 182182 three small nunataks (at 70°25–26′n, 51°58–59′w) in the ice cap 6–8 km south of umiartorfiup qaqqaa. mostly, only a few flows of dark basalt are seen together with traces of a sediment cover on the eroded gneiss (central nuussuaq section, panel j–k). lava flows and intercalated sediments. a c. 300 m thick section through the youngest volcanic and sedimentary lithologies is excellently exposed on a nunatak 3 km north of makittarissagaq (figs 163–166). the sediment locality is of difficult access and has only been investigated on the ground during a short helistop in 2016. four flows are preserved, of which the uppermost is a small erosional remnant. the two upper flows appear to be subaerial, whereas the two lower ones have the appearance of sills intruded into more than 50 m of quartzo-feldspathic clastic sediments. the lowermost sill is up to 100 m thick. the sills are almost certainly invasive lava flows. parts of the sediments are deformed and partially melted by the sills, but in the easternmost part of the exposure, there are more than 50 m of undeformed sediments (fig. 164). figure 165 presents an overview of the sedimentary succession. the lower invasive lava flow cuts the succession at a low angle, whereas the flow above the sediments appears to be parallel to the bedding. the sedimentary succession is c. 50 m thick and comprises six intervals, a to f (fig. 166). the lower sandstone (a) is cross-laminated, fine-grained and contains comminuted plant debris. the sandstone is interbedded with sand-streaked silty mudstone (b) with a few thin tuff beds. (c) is a channelised, cross-beddded, medium-grained sandstone. the overlying, coal-rich bed (d) is clayey, contains large pieces of coalified wood, and becomes sandier upwards. the sandstone (e) is similar to (a), and grades up into mudstone with tuff layers and coalified wood. the uppermost mudstone (f) is baked black towards the overlying lava. the succession is tentatively interpreted as lacustrine deposits, truncated by a minor fluvial channel (c). chemical compositions of the niaqussat member the large majority of niaqussat member rocks are tholeiitic basalts and picrites; only a few flows of basaltic andesite occur. the basaltic andesites, like those of the nordfjord member, have much higher contents of mgo, mf158ny_11cmv7pm5c6kmkb2cut_c.eps 520 532 inv 50 m f e d c b a fig. 166. sediments between invasive lava flows of the upper niaqussat member on eastern nuussuaq. the six intervals a to f are described in the text. the succession is tentatively interpreted as lacustrine deposits truncated by a minor fluvial channel. for location, see fig. 165. 183 ni, and cr than their orogenic namesakes and are very different rocks. representative chemical analyses of the niaqussat member rocks are shown in table 8, and all analyses are plotted in the variation diagrams, figs 167–169. as for the nordfjord member, mgo was chosen as the main variation parameter because the variations in feo* invalidate the mg-number as a common variation parameter. major elements the successively more evolved character (lower mgo) of the three units of the niaqussat member is clearly seen in fig. 167, where the separate identity of the middle unit is evident. the niaqussat member rocks are distinctly subalkaline, with total alkalis less than 3.4 wt%. of the basalts (mgo <12 wt%), 90 of 129 analysed samples (70%) are quartz normative. all basalts with mgo <8 wt% (59 samples) are quartz normative. the more silica-saturated character of the basalts and picrites compared with the rinks dal member is seen in the silica diagram of fig. 167, where the compositional field of the niaqussat member is situated at a higher level than that of the rinks dal member. the niaqussat member basalts with less than 10 wt% mgo have on average 50.4 wt% sio2 compared with 49.2 wt% sio2 in similar basalts of the rinks dal member. likewise, the niaqussat member picrites (with 12–15.2 wt% mgo) have on average 49.3 wt% sio2 compared with 47.7 wt% sio2 in similar uncontaminated picrites of the rinks dal member and the underlying vaigat formation. thus all the rocks of the niaqussat member have increased sio2 relative to uncontaminated rocks of the earlier parts of the volcanic succession. further in comparison with the rinks dal member, the niaqussat member basalts and picrites have generally higher al2o3 and lower tio2, feo*, na2o, and p2o5. the niaqussat member basalts generally have tio2 contents lower than 2 wt%. only in the upper niaqussat member are tio2 contents increased to 2–3 wt% (fig. 167); four flows have unusually high tio2 (3.9–5.2 wt%). two of these flows are the two uppermost niaqussat member flows in the sapernuvik profile, lying immediately beneath the flows of the sapernuvik member; the uppermost flow has the highest tio2 content in the entire maligât formation. the two other flows are the uppermost two of the seven flows of the upper niaqussat member in the point 2080 m profile on north-eastern nuussuaq, situated 130 km from sapernuvik. the four flows have many chemical characters in common, and it is therefore likely that the upper niaqussat member is practically complete at point 2080 m, despite the extensive erosion in the intervening area between the two profiles. the few basaltic andesites only cover a fairly narrow range in mgo (7.2–10 wt%). this is in strong contrast to the nordfjord member where the basaltic andesites range down to 3.4 wt% mgo and grade continuously into andesites and dacites. the difference suggests fewer high-level magma chambers and much shorter residence times in these for the niaqussat member magmas. trace elements variation diagrams for trace elements are shown in figs 168, 169. many elements show more scatter than in the rinks dal member, e.g. rb, ba, and ce. in general, the lower niaqussat member rocks have lowered sr, nb, v, cu and ni, and increased rb, ba and cr relative to the rinks dal member. the middle and upper niaqussat member basalts are in many respects quite similar to the rinks dal member basalts but have lower sr and higher sc. the basaltic andesites have higher rb, ba, zr, zr/y, ce and cr and lower v and zn than the basalts and picrites. ree and multi-element diagrams are shown in fig. 170. the picrites and basalts, including the evolved basalts of the upper niaqussat member, have lan/smn ratios close to but above 1. this clearly distinguishes the basalts of the upper niaqussat member from the otherwise quite similar basalts of the rinks dal member (fig. 100). the basaltic andesites have higher lan/smn ratios than other rocks of the member. in the multi-element diagrams the picrites and basalts all have nb-ta troughs and do not show the k and pb troughs that characterise the rinks dal member basalts, or even have k and pb peaks. the basaltic andesites show these features more pronouncedly and also have visible eu-ti troughs. the tirich basalt flows just below the sapernuvik member are different from all other rocks in the niaqussat member: they have no nb-ta troughs, they have k and pb troughs, and they appear not to be crustally contaminated. in conclusion, except for the latest high-ti flows all the rocks of the niaqussat member are to some degree crustally contaminated. processes and degrees of contamination are discussed in the last chapter. 184184 45 50 55 60 65 70 0 2 4 6 8 10 12 14 16 rinks dal mb 5 10 15 20 0 2 4 6 8 10 12 14 16 0 5 10 15 0 2 4 6 8 10 12 14 16 0 5 10 15 0 2 4 6 8 10 12 14 16 0 1 2 3 4 5 0 2 4 6 8 10 12 14 16 0 1 2 3 4 5 0 2 4 6 8 10 12 14 16 0 1 2 3 4 0 2 4 6 8 10 12 14 16 0.8 0.6 0.4 0.2 0.0 0 2 4 6 8 10 12 14 16 sapernuvik member (unit 540) upper niaqussat mb (unit 532) middle niaqussat mb (unit 531) basaltic andesite and andesite (unit 530) basalt and picrite, lower nia mb (unit 530) rinks dal mb rinks dal mb rinks dal mb rinks dal mb rinks dal mb rinks dal mb rinks dal mb sio2 al2o3 feo* cao tio2 na2o k2o p2o5 fig. 167. major-element variation diagrams for rocks of the niaqussat member. four data points for the sapernuvik member are included. data in wt% oxides recalculated volatile-free. feo* is total iron as feo. 185 0 10 20 30 40 50 60 70 80 90 100 0 2 4 6 8 10 12 14 16 0 100 200 300 400 500 600 0 2 4 6 8 10 12 14 16 0 100 200 300 400 0 2 4 6 8 10 12 14 16 0 10 20 30 40 50 60 70 80 0 2 4 6 8 10 12 14 16 0 100 200 300 400 0 2 4 6 8 10 12 14 16 0 5 10 15 20 25 30 35 40 0 2 4 6 8 10 12 14 16 0 1 2 3 4 5 6 7 8 9 10 0 2 4 6 8 10 12 14 16 0 10 20 30 40 50 60 70 80 0 2 4 6 8 10 12 14 16 rinks dal mb rinks dal mb rinks dal mb rinks dal mb rinks dal mb rinks dal mb rinks dal mb rinks dal mb sapernuvik member (unit 540) upper niaqussat mb (unit 532) middle niaqussat mb (unit 531) basaltic andesite and andesite (unit 530) basalt and picrite, lower nia. mb (unit 530) rb ba sr y mgo (wt%) zr mgo (wt%) nb zr/y ce fig. 168. incompatible trace-element variation diagrams for rocks of the niaqussat member. four data points for the sapernuvik member are included. data in ppm. 186186 vf vf 0 100 200 300 400 500 0 2 4 6 8 10 12 14 16 0 10 20 30 40 50 0 2 4 6 8 10 12 14 16 0 100 200 300 400 500 0 2 4 6 8 10 12 14 16 0 50 100 150 200 0 2 4 6 8 10 12 14 16 0 100 200 300 400 0 2 4 6 8 10 12 14 16 0 200 400 600 800 1000 0 2 4 6 8 10 12 14 16 rinks dal mb rinks dal mb rinks dal mb rinks dal mb rinks dal mb rinks dal mb sapernuvik member (unit 540) upper niaqussat mb (unit 532) middle niaqussat mb (unit 531) basaltic andesite and andesite (unit 530) basalt and picrite, lower nia. mb (unit 530) v sc cu zn mgo (wt%) ni mgo (wt%) cr fig. 169. transition-element variation diagrams for rocks of the niaqussat member. four data points for the sapernuvik member are included. data in ppm. 187 fig. 170. ree and multi-element diagrams for representative rocks of the niaqussat member. data for the sapernuvik member are included. left diagram, chondrite normalised; right diagram, primitive mantle normalised; normalisation factors from mcdonough & sun (1995). 1 10 100 1000 la ce pr nd sm eu gd tb dy ho er tm yb lu picrites (mgo ≥12 wt%), lower nq. mb (unit 530) 1 10 100 1000 rb ba th u nb ta k la ce pb pr sr nd p sm zr hf eu ti gd tb dy y yb 1 10 100 1000 la ce pr nd sm eu gd tb dy ho er tm yb lu basalts (mgo <12 wt%), lower nq. mb (unit 530) 1 10 100 1000 rb ba th u nb ta k la ce pb pr sr nd p sm zr hf eu ti gd tb dy y yb 1 10 100 1000 la ce pr nd sm eu gd tb dy ho er tm yb lu basaltic andesites, lower nq. mb (unit 530) 1 10 100 1000 rb ba th u nb ta k la ce pb pr sr nd p sm zr hf eu ti gd tb dy y yb 1 10 100 1000 la ce pr nd sm eu gd tb dy ho er tm yb lu sapernuvik member (unit 540) upper niaqussat mb (unit 532) middle niaqussat mb (unit 531) uppermost, high-ti flows 1 10 100 1000 rb ba th u nb ta k la ce pb pr sr nd p sm zr hf eu ti gd tb dy y yb 5301 5301 5301 5301 5301 5301 5301 5301 5301 5301 5301 5301 332859 263973 156625 176642 340824 176603 264057 274415 318827 328420 362179 156571 5219.82 5450.71 5445.76 5441.90 5240.88 5455.12 5427.73 5442.51 5220.01 5358.23 5240.45 5445.94 7029.213 6948.649 7004.492 6957.254 6944.488 6940.704 6952.778 7010.641 6946.544 6923.912 6950.978 7004.470 1694 581 236 572 1162 312 1274 278 1193 872 1319 216 46.85 46.14 47.42 46.68 47.65 49.09 48.71 49.56 50.07 49.39 55.30 53.37 1.14 1.25 1.33 1.34 1.34 1.51 1.60 1.45 1.61 1.94 1.28 1.50 12.15 11.87 13.34 12.79 13.46 14.21 14.68 15.25 14.89 14.87 15.18 16.74 1.68 2.61 2.95 1.45 3.03 2.63 3.09 4.59 3.91 5.77 0.40 0.89 8.55 8.21 7.28 8.61 7.62 8.05 7.45 5.96 7.13 6.40 7.50 7.49 0.16 0.17 0.16 0.20 0.17 0.18 0.18 0.17 0.18 0.19 0.14 0.14 14.53 14.50 13.12 12.04 11.61 9.03 8.05 7.77 7.61 7.23 9.13 7.64 9.15 9.77 8.21 9.96 9.44 11.29 11.25 11.45 11.33 10.73 7.41 7.64 1.26 1.01 1.30 1.00 1.51 1.97 1.79 2.00 2.12 2.11 2.04 1.94 0.178 0.280 0.120 0.240 0.170 0.328 0.120 0.208 0.312 0.379 0.442 0.640 0.099 0.100 0.100 0.070 0.116 0.133 0.140 0.142 0.152 0.167 0.172 0.160 4.36 4.30 3.94 5.48 3.82 1.53 2.66 1.80 0.79 1.11 1.20 1.31 100.11 100.21 99.27 99.86 99.93 99.95 99.72 100.35 100.10 100.28 100.19 99.46 10.06 10.56 9.93 9.91 10.35 10.42 10.23 10.09 10.65 11.59 7.86 8.29 74.50 73.53 72.76 71.07 69.42 63.68 61.41 60.90 59.11 55.78 70.14 65.08 72.2 77.7 77.1 76.2 81.6 78.2 82.2 75.8 86.2 92.7 69.8 84.3 81.8 71.4 48.0 38.3 97.7 95.0 94.6 60.9 124 119 38.4 30.7 413 465 122 151 208 171 123 91.5 107 95.9 113 36.7 37.7 32.5 32.3 32.8 33.5 38.8 40.0 37.6 40.1 33.1 28.6 30.9 278 261 272 267 274 301 308 313 319 323 215 253 1401 1037 1055 1114 856 527 404 651 392 293 715 523 12.3 12.6 17.2 13.2 17.1 20.0 18.5 18.5 19.7 20.9 19.9 22.8 4.78 4.68 1.77 2.68 1.88 4.64 0.79 1.82 4.44 6.50 13.5 15.8 242 144 129 137 135 188 145 180 173 210 206 231 19.0 21.6 21.6 20.9 22.2 25.3 28.2 25.7 30.3 28.8 24.0 27.0 82.8 72.7 91.6 88.4 91.4 93.3 94.8 91.7 99.9 114 175 168 3.35 2.85 3.69 3.57 3.78 3.47 3.78 3.94 3.98 5.15 7.48 8.09 0.067 0.047 0.044 0.048 0.150 0.133 0.030 0.032 0.042 0.091 0.368 0.415 55.3 36.0 57.0 69.3 53.5 48.0 47.1 56.3 67.5 54.2 229 262 5.34 4.07 6.60 6.28 5.48 4.88 5.72 5.54 6.19 6.67 17.6 21.0 13.3 10.6 16.0 15.5 14.5 13.3 14.4 13.9 15.2 17.2 39.8 44.7 2.04 1.61 2.30 2.26 2.11 2.05 2.27 2.17 2.41 2.82 4.81 5.45 9.78 8.24 10.6 10.6 10.7 10.6 11.4 10.8 12.0 14.0 20.0 22.3 2.86 2.54 3.04 3.00 3.20 3.40 3.43 3.21 3.75 4.17 4.60 4.90 0.984 0.943 1.04 1.05 1.11 1.22 1.20 1.24 1.31 1.50 1.28 1.42 3.47 3.13 3.69 3.62 3.88 4.23 4.18 3.86 4.47 4.89 5.12 5.05 0.592 0.545 0.614 0.631 0.648 0.723 0.743 0.684 0.765 0.816 0.762 0.832 3.50 3.38 3.73 3.71 3.89 4.39 4.55 4.07 4.81 4.90 4.42 4.74 0.710 0.704 0.798 0.800 0.805 0.906 0.951 0.857 0.985 0.994 0.856 0.984 1.93 1.90 2.03 2.02 2.18 2.46 2.57 2.37 2.64 2.59 2.33 2.52 0.282 0.284 0.296 0.295 0.315 0.360 0.374 0.352 0.394 0.369 0.340 0.367 1.69 1.70 1.82 1.86 1.93 2.21 2.34 2.09 2.46 2.27 2.04 2.26 0.241 0.258 0.268 0.262 0.276 0.309 0.350 0.310 0.359 0.328 0.293 0.337 2.15 1.89 2.44 2.43 2.41 2.48 2.58 2.45 2.60 2.99 4.28 4.37 0.212 0.194 0.268 0.274 0.242 0.216 0.250 0.260 0.286 0.322 0.552 0.596 1.36 0.927 1.75 1.40 1.50 0.999 1.28 1.31 1.18 1.02 4.38 8.80 0.896 0.511 1.16 1.03 0.816 0.629 0.834 0.774 0.817 0.664 4.03 4.39 0.220 0.127 0.216 0.256 0.174 0.173 0.152 0.194 0.191 0.185 1.03 1.05 0.705004 0.711708 6.43 101.59 0.512814 0.511796 4.93 -14.93 17.149 16.950 15.094 15.038 37.346 37.270 lith. code ggu no. deg. w deg. n altitude, m sio2 tio2 al2o3 fe2o3 feo mno mgo cao na2o k2o p2o5 volatiles sum feo* mg number zn cu ni sc v cr ga rb sr y zr nb cs ba la ce pr nd sm eu gd tb dy ho er tm yb lu hf ta pb th u isotope ratios calculated at 60 ma 87sr/86sr60 εsr 143nd/144nd60 εnd 206pb/204pb60 207pb/204pb60 208pb/204pb60 for explanation of lithological codes, see table 1. for petrographical notes on the samples, see table 8c. geographical coordinates in wgs 84. first two digits are degrees, then follow minutes in decimal form. major elements in wt% (xrf analyses). trace elements in ppm (zn–ga: xrf analyses; rb–u: icp-ms analyses). feo* = total iron as feo. mg number = 100 × atomic mg/(mg+fe2+), with the iron oxidation ratio adjusted to fe2o3/feo = 0.15. table 8a. chemical analyses of rocks of the niaqussat member picrite basalt basaltic andesitelithology lower niaqussat member 189 5311 5311 5321 5321 5321 5321 6181 6181 6281 6281 176622 274419 332898 340817 332894 176615 318753 318707 318824 156570 5455.19 5442.48 5226.30 5240.27 5226.51 5455.73 5227.79 5217.96 5209.12 5445.91 6940.777 7010.613 7030.746 6944.813 7030.736 6941.272 7013.735 7004.028 6949.517 7004.442 361 323 1897 1306 1977 543 1272 371 10 175 50.41 49.00 49.08 50.44 48.38 45.10 49.43 49.45 48.23 49.13 1.78 1.71 2.63 2.69 3.82 5.11 1.60 1.45 1.35 1.41 14.07 15.68 12.93 12.73 12.62 11.91 14.60 13.73 13.24 13.80 2.89 6.57 6.58 4.64 4.64 6.69 1.51 1.47 2.14 2.48 9.54 4.81 8.02 9.80 10.34 10.97 9.42 9.07 8.45 7.66 0.20 0.18 0.22 0.24 0.23 0.29 0.19 0.17 0.17 0.16 6.45 6.10 5.91 5.07 5.54 5.36 7.60 11.37 12.64 11.15 11.04 11.78 10.30 9.42 9.67 9.73 12.04 9.82 10.05 9.37 2.37 1.98 2.44 2.42 2.40 2.05 1.94 1.81 1.63 1.76 0.182 0.113 0.267 0.647 0.630 0.326 0.180 0.365 0.139 0.390 0.163 0.156 0.268 0.286 0.426 0.646 0.160 0.152 0.150 0.110 0.90 2.51 0.96 1.22 0.80 1.38 1.25 1.01 1.39 1.83 100.00 100.59 99.60 99.59 99.49 99.56 99.92 99.87 99.58 99.25 12.14 10.72 13.94 13.98 14.52 16.99 10.78 10.39 10.38 9.89 51.80 53.50 46.16 42.32 43.57 38.95 58.78 68.88 71.13 69.51 94.2 89.7 120 134 138 169 83.3 87.4 83.8 80.3 91.9 90.3 259 224 323 444 131 78.2 98.7 62.8 50.9 108 56.6 47.0 75.3 47.0 126 214 319 98.31 43.6 39.0 49.5 42.8 39.9 42.6 40.7 32.6 33.4 32.4 342 313 479 354 442 510 325 288 291 280 137 375 103 67.6 176 74.5 409 851 926 1151 19.8 20.8 21.8 22.4 22.5 23.6 19.4 18.9 17.4 17.4 3.88 0.75 7.57 23.3 12.3 1.88 4.99 8.90 1.61 11.1 155 184 207 156 207 187 169 185 191 159 33.0 31.2 43.9 53.2 56.5 78.2 28.9 23.4 22.2 21.4 114 110 185 202 276 411 97.3 104 89.8 98.2 4.55 4.37 9.93 8.88 21.4 33.9 3.80 4.58 4.06 4.11 0.118 0.018 0.175 0.590 0.146 0.017 0.126 0.253 1.01 0.450 67.4 35.9 84.6 120 154 184 46.6 86.2 48.8 79.8 6.92 6.59 12.7 13.2 17.8 27.3 5.68 7.41 6.00 7.21 17.4 16.4 30.9 31.7 43.1 63.7 14.8 18.7 15.5 17.9 2.62 2.63 4.53 4.61 6.35 9.17 2.28 2.78 2.29 2.44 13.3 13.0 21.9 23.0 30.7 44.2 11.8 13.4 11.3 11.6 4.20 4.10 6.50 6.90 8.81 12.1 3.71 3.64 3.22 3.30 1.43 1.40 2.09 2.08 2.60 3.49 1.30 1.26 1.17 1.14 5.19 4.88 7.78 8.48 10.4 14.4 4.64 4.27 3.78 3.90 0.924 0.853 1.30 1.41 1.71 2.36 0.785 0.728 0.657 0.656 5.67 5.10 7.72 8.77 10.1 13.8 5.14 4.03 3.77 3.89 1.18 1.08 1.56 1.85 2.01 2.76 1.04 0.820 0.767 0.793 3.30 2.86 4.36 5.04 5.50 7.56 2.79 2.19 1.98 2.12 0.476 0.423 0.643 0.755 0.781 1.09 0.412 0.322 0.297 0.311 3.01 2.53 3.91 4.60 4.85 6.71 2.48 1.93 1.86 1.91 0.429 0.372 0.550 0.655 0.694 0.943 0.374 0.290 0.266 0.267 3.05 2.98 4.68 5.23 6.72 9.63 2.62 2.72 2.31 2.66 0.287 0.276 0.560 0.538 1.22 1.88 0.266 0.355 0.280 0.360 1.97 1.36 2.02 3.08 2.11 2.96 1.17 1.87 1.49 1.75 1.05 0.963 1.49 1.75 2.08 3.18 0.743 1.13 0.828 1.22 0.324 0.272 0.437 0.569 0.613 0.844 0.218 0.333 0.239 0.317 0.704381 0.703377 0.703396 -2.40 -16.65 -16.38 0.512710 0.512883 0.512869 2.92 6.29 6.02 17.300 18.042 18.151 15.142 15.370 15.361 37.382 37.792 37.882 lith. code ggu no. deg. w deg. n altitude, m sio2 tio2 al2o3 fe2o3 feo mno mgo cao na2o k2o p2o5 volatiles sum feo* mg number zn cu ni sc v cr ga rb sr y zr nb cs ba la ce pr nd sm eu gd tb dy ho er tm yb lu hf ta pb th u isotope ratios calculated at 60 ma 87sr/86sr60 εsr 143nd/144nd60 εnd 206pb/204pb60 207pb/204pb60 208pb/204pb60 table 8b. chemical analyses of rocks of the niaqussat member and possible feeder dykes ti-rich top flowsbasalt basalt weakly contaminated basalt dykeslithology middle niaqussat member upper niaqussat member possible feeders niaqussat mb feeder 190190 table 8c. notes on analysed samples of the niaqussat member and possible feeder dykes 332859. picrite with well-crystallised doleritic groundmass and abundant microphenocrysts of olivine with tiny enclosed chromites. zeolites and carbonate in vugs and groundmass. lowest flow in the niaqussat member, nunavik profile, eastern nuussuaq. 263973. picrite with well-crystallised doleritic groundmass and abundant up to 1 mm olivine phenocrysts with tiny enclosed chromites. zeolites in vugs and groundmass. lava flow, ikorfarsuit profile, mellemfjord, west disko. 156625. picrite with olivine microphenocrysts. lava flow 2 m thick, north wall of rink dal c. 1.5 km from the coast, north-west disko. 176642. picrite close to magnesian basalt with well-crystallised doleritic groundmass and up to 1 mm large olivine phenocrysts with enclosed chromite crystals. lava flow 3 m thick, kingittuusaq profile, outer nordfjord, west disko. 340824. picrite close to magnesian basalt with well-crystallised doleritic groundmass and scattered up to 1 mm olivine phenocrysts (mostly pseudomorphosed). rare up to 3 mm glomerocrysts of plagioclase and pseudomorphosed olivine. traces of tiny chromites. lava flow 5 m thick, frederik lange dal profile, east disko. 176603. basalt with well-crystallised doleritic groundmass and scattered up to 1 mm phenocrysts of olivine. lava flow 10 m thick, sapernuvik profile, west disko. 264057. basalt with well-crystallised groundmass and common up to 2 mm phenocrysts and glomerocrysts of plagioclase and scarce up to 1 mm olivine penocrysts (pseudomorphosed). lava flow 12 m thick, point 1300 m profile, vesterdalen, west disko. 274415. almost aphyric basalt with very fine-grained groundmass and very scarce microphenocrysts of plagioclase and olivine (pseudomorphosed). lava flow 14 m thick, ‘point 440 m northern gully’ profile, north of hammer dal, north-west disko. 318827. basalt with very fine-grained groundmass and scarce <1 mm glomerocrysts of plagioclase and olivine (pseudomorphosed). uppermost exposed lava flow on plateau, kvandalen profile, east disko. 328420. basalt with very fine-grained groundmass and widespread phenocrysts and up to 4 mm glomerocrysts of plagioclase, olivine (partly pseudomorphosed) and augite. lava flow 12 m thick, alanngup qaqqai profile, south-west disko. 362179. magnesian basaltic andesite with rare up to 1 cm bodies of native iron rimmed by troilite. abundant up to 1 mm phenocrysts and micro phenocrysts of orthopyroxene and scarce up to 1 mm phenocrysts of olivine, sometimes with chromite. widespread xenocrysts and micro xenoliths of plagioclase and red spinel, and rare xenocrysts of quartz. lava flow 5 m thick, qinngusaq profile, inner kvandalen, east disko. 156571. basaltic andesite; thin section lost. lava flow, north wall of rink dal c. 1.5 km from the coast, north-west disko. 176622. basalt with fine-grained groundmass and common phenocrysts and glomerocrysts of plagioclase up to 1.5 mm large and very scarce microphenocrystic augite. lava flow 8 m thick, sapernuvik profile, west disko. 274419. basalt with well-crystallised doleritic groundmass, common up to 2 mm glomerocrysts of plagioclase and scarce microphenocrysts of olivine (pseudomorphosed) and augite. lava flow 8 m thick, ‘point 440 m northern gully’ profile, north of hammer dal, north-west disko. 332898. basalt with well-crystallised fine-grained groundmass rich in equidimensional fe-ti oxide grains. scattered microphenocrystic plagioclase and augite. glomerocrysts <1.5 mm of plagioclase and augite. entablature-dominated lava flow 20 m thick, point 2080 m profile, eastern nuussuaq. 340817. basalt; thin section lost. lava flow 15 m thick, frederik lange dal profile, east disko. 332894. evolved high-ti basalt; thin section lost. lava flow 20 m thick, point 2080 m profile, eastern nuussuaq. 176615. evolved high-ti basalt with fine-grained groundmass rich in fe-ti oxide grains. common phenocrytsts of plagioclase, olivine and augite. there are no fe-ti oxide microphenocrysts. most ti-rich basalt in the maligât formation. lava flow 5 m thick, sapernuvik profile, west disko. 318753. picrite dyke. glass chill with abundant microphenocrystic olivine and plagioclase up to 1 mm large. the crystalline groundmass is composed of plagioclase, clinopyroxene, olivine, ilmenite and residuum, and there is abundant microphenocrystic olivine and plagioclase. dyke 1 m thick cutting lava flows of the rinks dal member. giesecke monument, south coast of nuussuaq. 318707. magnesian basalt dyke with well-crystallised doleritic groundmass of plagioclase, clinopyroxene, olivine, fe-ti oxides and residuum and scattered olivine phenocrysts up to 1.5 mm and abundant olivine microphenocrysts <0.5 mm large. dyke 3–4 m thick, striking ne–sw, cutting sediments of the atane formation east of atanikerluk, south nuussuaq. 318824. picrite dyke. the groundmass is well crystallised and composed of plagioclase with characteristic leaf-textured hollow-cored grains, clinopyroxene, olivine, fe-ti oxide and residuum. the abundant microphenocrystic olivines (<0.5 mm) are much smaller than the plagioclase grains, and the rock does not therefore appear as porphyritic. dyke 3.5 m thick, striking 60°ne, shore north of inngigissoq, east disko. 156570. magnesian basalt feeder dyke for a niaqussat member lava flow. abundant up to 1 mm phenocrysts and microphenocrysts of olivine, a few with chromite. the doleritic groundmass is composed of plagioclase, lowand high ca clinopyroxene, ilmenite and residuum. north wall of rink dal c. 1.5 km from the coast, north-west disko. phenocryst phases as observed in thin section are mentioned in order of decreasing abundance. samples of subaerial lava flows are usually taken in massive columns a few metres above the flow base. 191 5401 5401 5401 5401 176611 176614 176608 176609 5454.77 5455.80 5455.74 5455.48 6942.010 6941.302 6941.734 6941.725 637.29 559.41 584.05 596.39 47.37 47.85 49.01 48.96 0.80 1.11 1.54 1.50 16.09 14.69 14.75 14.68 2.57 3.96 3.84 6.04 6.70 6.92 8.00 5.89 0.15 0.19 0.19 0.19 10.59 9.80 7.50 7.37 12.71 11.21 11.72 12.02 1.57 1.71 2.22 2.21 0.188 0.126 0.248 0.258 0.062 0.099 0.171 0.171 1.75 2.48 1.33 1.04 100.55 100.15 100.52 100.33 9.01 10.48 11.46 11.32 70.39 65.41 56.98 56.83 54.8 73.2 80.1 83.6 118 227 161 151 234 189 91 93 38.6 44.7 35.2 39.9 229 270 312 298 440 465 198 214 51.7 50.5 48.4 48.8 14.8 15.8 18.5 18.9 4.82 1.09 1.86 2.93 142 101 176 178 15.7 23.8 26.8 27.6 40.4 60.0 83.2 84.2 1.48 4.16 7.77 8.10 0.009 0.153 0.009 0.015 17.7 35.7 74.6 70.8 1.82 3.33 6.41 6.64 5.20 8.82 16.0 15.7 0.850 1.27 2.42 2.41 4.87 6.74 11.5 11.5 1.74 2.44 3.40 3.35 0.699 0.936 1.25 1.23 2.22 3.22 4.26 4.16 0.422 0.617 0.736 0.713 2.69 3.91 4.59 4.59 0.585 0.869 0.986 0.957 1.64 2.53 2.63 2.55 0.238 0.365 0.395 0.400 1.49 2.35 2.41 2.39 0.211 0.346 0.367 0.368 1.16 1.72 2.18 2.33 0.194 0.244 0.484 0.317 0.417 0.565 0.680 0.143 0.331 0.620 0.636 0.043 0.094 0.143 0.167 0.703382 0.703306 -16.58 -17.66 0.512971 0.512929 7.99 7.18 17.669 17.807 15.304 15.431 37.602 37.608 table 9a. chemical analyses of rocks of the sapernuvik member lithology basalt lith. code ggu no. deg. w deg. n altitude, m sio2 tio2 al2o3 fe2o3 feo mno mgo cao na2o k2o p2o5 volatiles sum feo* mg number zn cu ni sc v cr co ga rb sr y zr nb cs ba la ce pr nd sm eu gd tb dy ho er tm yb lu hf ta pb th u isotope ratios calculated at 60 ma 87sr/86sr60 εsr 143nd/144nd60 εnd 206pb/204pb60 207pb/204pb60 208pb/204pb60 for explanation of lithological codes, see table 1. geographical coordinates in wgs 84. first two digits are degrees, then follow minutes in decimal form. major elements in wt% (xrf analyses). trace elements in ppm (zn–ga: xrf analyses; rb–u: icp-ms analyses). feo* = total iron as feo. mg number = 100 × atomic mg/(mg+fe2+), with the iron oxidation ratio adjusted to fe2o3/feo = 0.15. 5401 5401 5401 5401 176611 176614 176608 176609 5454.77 5455.80 5455.74 5455.48 6942.010 6941.302 6941.734 6941.725 637.29 559.41 584.05 596.39 47.37 47.85 49.01 48.96 0.80 1.11 1.54 1.50 16.09 14.69 14.75 14.68 2.57 3.96 3.84 6.04 6.70 6.92 8.00 5.89 0.15 0.19 0.19 0.19 10.59 9.80 7.50 7.37 12.71 11.21 11.72 12.02 1.57 1.71 2.22 2.21 0.188 0.126 0.248 0.258 0.062 0.099 0.171 0.171 1.75 2.48 1.33 1.04 100.55 100.15 100.52 100.33 9.01 10.48 11.46 11.32 70.39 65.41 56.98 56.83 54.8 73.2 80.1 83.6 118 227 161 151 234 189 91 93 38.6 44.7 35.2 39.9 229 270 312 298 440 465 198 214 51.7 50.5 48.4 48.8 14.8 15.8 18.5 18.9 4.82 1.09 1.86 2.93 142 101 176 178 15.7 23.8 26.8 27.6 40.4 60.0 83.2 84.2 1.48 4.16 7.77 8.10 0.009 0.153 0.009 0.015 17.7 35.7 74.6 70.8 1.82 3.33 6.41 6.64 5.20 8.82 16.0 15.7 0.850 1.27 2.42 2.41 4.87 6.74 11.5 11.5 1.74 2.44 3.40 3.35 0.699 0.936 1.25 1.23 2.22 3.22 4.26 4.16 0.422 0.617 0.736 0.713 2.69 3.91 4.59 4.59 0.585 0.869 0.986 0.957 1.64 2.53 2.63 2.55 0.238 0.365 0.395 0.400 1.49 2.35 2.41 2.39 0.211 0.346 0.367 0.368 1.16 1.72 2.18 2.33 0.194 0.244 0.484 0.317 0.417 0.565 0.680 0.143 0.331 0.620 0.636 0.043 0.094 0.143 0.167 0.703382 0.703306 -16.58 -17.66 0.512971 0.512929 7.99 7.18 17.669 17.807 15.304 15.431 37.602 37.608 table 9a. chemical analyses of rocks of the sapernuvik member lithology basalt lith. code ggu no. deg. w deg. n altitude, m sio2 tio2 al2o3 fe2o3 feo mno mgo cao na2o k2o p2o5 volatiles sum feo* mg number zn cu ni sc v cr co ga rb sr y zr nb cs ba la ce pr nd sm eu gd tb dy ho er tm yb lu hf ta pb th u isotope ratios calculated at 60 ma 87sr/86sr60 εsr 143nd/144nd60 εnd 206pb/204pb60 207pb/204pb60 208pb/204pb60 for explanation of lithological codes, see table 1. geographical coordinates in wgs 84. first two digits are degrees, then follow minutes in decimal form. major elements in wt% (xrf analyses). trace elements in ppm (zn–ga: xrf analyses; rb–u: icp-ms analyses). feo* = total iron as feo. mg number = 100 × atomic mg/(mg+fe2+), with the iron oxidation ratio adjusted to fe2o3/feo = 0.15. 176611 fine-grained basalt with up to 3 mm large phenocrysts of plagioclase and up to 1.5 mm olivine, and up to 3 mm glomerocrysts of plagioclase with resorbed outlines. third and youngest preserved flow in the sapernuvik member; sapernuvik profile, south-west disko. 176614 fine-grained basalt with common 1–2 mm large phenocrysts of plagioclase, olivine and augite, and scattered up to 1 cm large, gabbroic inclusions with augite and plagioclase. lowest flow in the sapernuvik member, sapernuvik profile, south-west disko. 176608 fine-grained basalt with up to 2 mm large glomerocrysts and phenocrysts of plagioclase and olivine (pseudomorphed) and <1 mm large augite. second flow in the sapernuvik member, sapernuvik profile, south-west disko. 176609 fine-grained basalt with sparse phenocrysts and up to 3 mm large glomerocrysts of plagioclase, augite and olivine (pseudomorphed). central part of second flow in the sapernuvik member, sapernuvik profile, south-west disko. phenocryst phases as observed in thin section are mentioned in order of decreasing abundance. samples of subaerial lava flows were normally collected in massive columns a few metres above the flow base. table 9b. notes on analysed samples of the sapernuvik member footnotes, table 9 192192 sapernuvik member summary of the main features of the sapernuvik member • the youngest lava flows on disko and nuussuaq east of the itilli fault. • only three lava flows preserved in an erosional remnant on southern west disko; original thickness and extent unknown. • conformably overlies lava flows of the upper niaqussat member. • uncontaminated, relatively magnesian basalts. one flow may include a component of remelted gabbro cumulates. lithostratigraphy of the sapernuvik member new member history. the sapernuvik member lava succession is shown on the geological map 1:100 000 mellemfjord 69 v.1 nord as unit βol (pedersen & ulff-møller 1987) and on the geological section along the south coast of disko (sapernuvik unit, ni5; pedersen et al. 2003). the flows were briefly described by pedersen (1977b, p. 47–48 and table 8 nos 11–13). name. after sapernuvik, a small cove on the west coast of disko (fig. 4). distribution. the sapernuvik member is only preserved within an area measuring 4 × 6 km in a down-thrown, west-dipping fault block on southern west disko. type section. sapernuvik on southern west disko, on the mountaintop above 560 m (fig. 14, profile 1; figs 161 (photograph) and 171 (south disko section at 0–5 km)). thickness. 80 m. lithology. three subaerial lava flows of olivine-microphyric basalt. subdivisions. none. boundaries. the sapernuvik member conformably overlies the lava flows of the niaqussat member. the upper boundary is erosional. age. inferred to be paleocene, 61–60 ma, magnetochron c26r, based on radiometric dating (storey et al. 1998) of one of the youngest flows in the underlying niaqussat member. correlation. the sapernuvik member is not known to correlate with any other lava succession in the region. 1 2 3 5 km 1000 m 200 400 600 800 sapernuvikpaakkarut gv03_03_061_lml 540 532 531? 530 520 518 nw se fig. 171. the type section for the sapernuvik member (unit 540) and the upper niaqussat member (unit 532), and reference section for the lower and middle niaqussat member (units 530 and 531, respectively). photogrammetrically measured section of the coastal cliff on southwestern disko between paakkarut and sapernuvik. blue three-digit numbers are lithological codes. slightly modified excerpt from the south disko section (pedersen et al. 2003). see also photograph in fig. 161. 193 geology and geochemistry of the sapernuvik member the sapernuvik member (unit 540) comprises only a few lava flows. these are the youngest lava flows preserved on disko, and they are so compositionally different from all the other flows in the maligât formation that they have been assigned the status of a separate member. the lava flows of the sapernuvik member cap four separate flat mountaintops on westernmost disko south of mellemfjord between kingittup qaqqaa and sapernuvik (figs. 4, 161). the member has only been sampled in the type section. here, three flows are preserved capping a lava succession that dips 2.5–3°w (fig. 171). they conformably overlie flows of the upper niaqussat member (fig. 172) and are separated from these by a few decimetres of lateritic claystone. one or two more additional flows may be present on the top plateau of kingittup qaqqaa. the member is cut by several basalt dykes, which may have been feeders to inferred lava flows of the younger svartenhuk and naqerloq formations (larsen et al. 2016). lithologically, the dark grey lavas are fine-grained with phenocrysts of olivine and plagioclase up to a few millimetres in size. one flow also has up to 5 mm large augite phenocrysts. two of the flows have inclusions of dolerite and gabbro up to a few centimetres in size (fig. 173). the four existing analyses of lava flows of the sapernuvik member are given in table 9. they have been plotted with the niaqussat member in figs 167–170. the flows are distinctly more magnesian (7.5–10.7 wt% mgo) than the underlying evolved flows of the upper niaqussat member. they also have lower sio2 and higher cao and appear not to be crustally contaminated. the highest flow has very high al2o3 (16.3 wt%) and cao (12.9 wt%) without being plagioclase-accumulative; it also has low feo* and very low tio2 (0.8 wt%) and p2o5 (0.06 wt%). the sapernuvik member lavas have low contents of several incompatible trace elements and cr, and highish sc and cu contents. the ree patterns have gd/lu ratios lower than any other flows in the maligât formation, mf164_17cma1974_sapernuvik2no12amod2_c.eps 532 (176616) 540 (176614) 176608 540 (176614) 532 fig. 172. the two lowest lava flows of the sapernuvik member (540) conformably overlying lava flows of the upper niaqussat member (532). the same corrie between sapernuvik and kingittup qaqqaa as shown in fig. 162 but at another place. the numbered samples are taken in the indicated flows but at another locality; the actual sample profile is shown in fig. 161. 194194 suggesting melt formation at relatively shallow levels. the aland ca-rich flow (sample 176611) has relatively high rb, k, and sr contents in accordance with its high normative feldspar content (52%). larsen & pedersen (2009) suggested that the sapernuvik member lavas include a component of remelted gabbroic cumulates. mf165_8cm176614upcut_c.eps ol 5 mm cpx pl fig. 173. thin section of sparsely olivine-plagioclase phyric basalt of the sapernuvik member with a gabbro xenolith 1 cm large, consisting of augite (cpx) and plagioclase (pl). sample 176614, sapernuvik, south-west disko. 195 dyke systems of the nordfjord and niaqussat members summary of the main features of the contaminated dyke systems • weakly contaminated small dykes scattered over disko and nuussuaq. • strongly contaminated large dykes in three major zones on disko; structurally related to the disko gneiss ridge. • composite basalt/andesite dykes up to 80 km long, many with native iron and troilite. • magma chambers concentrated on north-west and east disko. • subvolcanic magma chambers and tube-shaped intrusions with native iron and sulfide accumulations. • economic geology: accumulations of native iron and sulfides have been subjected to exploration for ni and platinum-group elements. the dyke systems for the nordfjord and niaqussat members are treated together here because they are difficult to distinguish from each other and appear to have followed similar patterns. weakly contaminated dykes the basalts and picrites of the nordfjord and niaqussat members were presumably erupted mainly, but not solely, west of the disko gneiss ridge. on eastern disko and south-eastern nuussuaq there are scattered occurrences of dykes of silica-enriched magnesian basalt with 50–52 wt% sio2 and 7.5–13 wt% mgo, very similar to the less contaminated rocks of the nordfjord and niaqussat members, indicating that these magmas were also erupted over larger areas east of the disko gneiss ridge (fig. 174; larsen & pedersen 1992). similar dykes occur on western disko as single dykes or part of the larger, commonly composite intrusive systems. main dyke systems of the strongly contaminated magmas a considerable number of strongly sediment-contaminated dykes, craters and necks related to the nordfjord and niaqussat members have been located on disko. in contrast, no dykes of this kind are known from nuussuaq. many of these rocks carry native iron and sulfides. except for two dykes at the coast which were noted in the 1870s, these dykes were first located during the regional geological mapping, and later, as their economic potential became evident (pedersen 1975c), during systematic prospecting by mining companies (ulff-møller 1991). these intrusive systems, in which dykes, feeder dykes, necks and craters are associated, have given rise to spectacular localities with native iron, sulfides and sediment xenoliths. almost all the intrusions carry xenoliths of magma-modified mudstone and sandstone in various stages of melting and equilibration, indicating that the magmas were emplaced from high-level reservoirs situated within successions of mudstone and sandstone. many of the dykes are composite, and eruption through such feeders would have formed the composite lava flows found within both the nordfjord and niaqussat members. the dyke systems are shown in fig. 174. the intrusions have a characteristic distribution pattern apparently influenced by the disko gneiss ridge that extends s–n for about 80 km from the central south coast of disko to stordal in the north and continues northwards below exposure level to the kuugannguaq valley but not farther (chalmers et al. 1999). a number of intrusive systems can be recognised. system a comprises at least 10–12 composite dykes striking nw–se to nnw–sse within a zone c. 15 km wide between the central part of hammer dal and morten porsild dal on northern disko. several craters and volcanic necks are associated with the dyke bodies. the western part of the zone is poorly defined because of the strongly faulted and very poorly exposed areas along the northern west coast of disko. particularly important intrusions within system a, described below, are the hammer dal, hanekammen and nordfjord complexes and the inner giesecke dal intrusion. system b comprises at least two n–s-striking composite dykes, which extend for more than 50 km along the crest of the disko gneiss ridge between stordal in the north and the south coast of disko. one of the dykes, the killiit dyke, extends for more than 85 km. it originates as a nw–se-striking dyke within system a north-west of stordal; south of stordal it changes strike to n–s and continues for more than 70 km to the south coast of disko at killiit. system c comprises a few dykes striking nw–se within a 70 km long zone along the north-east coast of disko, 196196 gv01_02_162_lml.eps a c b 20 km 53°w54°w55°w maligât formation vaigât formation cretaceous and paleocene sediments precambrian gneiss and amphibolite fault crater or volcanic neck segments of large, strongly contaminated dykes small dyke, presumed niaqussat member large native iron boulders at uiffaq and in stordal 69°30'n 70°n maniillat dyke niaqussat hammer dal rink dal nordfjord mellemfjord h d c h kcn c sto rd al mp d al illukunnguaq dyke killiit dyke qeqertarsuaq uifffaq dyke uifffaq kvandalen kuugannguaq skansen 52°w eocene volcanic rocks and intrusions ice lake giesecke dal kangerluk rink’s dyke fig. 174. geological map of disko and southern nuussuaq showing the extension of the three main systems (green outlines a, b, c) of strongly contaminated, commonly native-ironand sulfide-bearing dykes and craters described in the text. some specific occurrences are indicated as explained in the legend. hdc: hammer dal complex. hkc: hanekammen complex. nc: nordfjord complex. mp dal: morten porsild dal. the black frame on north-west disko shows the location of the map in fig. 181. 197 semi-parallel to the vaigat strait. the easternmost dyke exposure is a sulfide-mineralised dyke at illukunnguaq. it is possible, but not proved, that a large composite, nativeiron-bearing lava flow in the kvandalen area could have been erupted from a system c dyke. others comprise a number of small dyke segments striking in various directions from n–s to nne–ssw to e–w, found between the uiffaq peninsula west of the disko gneiss ridge in the south and giesecke dal in the north. parts of the areas have not been searched in any detail and are likely to contain many dyke segments, and other areas, such as the outer west coast between nordfjord and giesecke dal, are extensively faulted and poorly exposed. there must also be hidden dykes in the shallow sea offshore western disko. strongly contaminated dykes and lavas are conspicuously absent in the area east of the disko gneiss ridge and south of kvandalen on eastern disko. evidently, the disko gneiss ridge had a confining influence which prevented horizontal flow of magma from the high-level magma reservoirs eastwards via dykes across the ridge. as seen for system b, the stress pattern along the crest of the gneiss ridge and within the overlying older basalts may have forced the magma conduits to change direction from se-bound to s-bound over distances of more than 50 km. only north of the disko gneiss ridge (north-east of the kuugannguaq valley), nw–se-trending dyke intrusions continued eastwards within system c, allowing strongly contaminated lavas and dykes to be formed on eastern disko. the strongly contaminated intrusions yield important information on the sediment–magma reactions in the high-level magma reservoirs and on the sediments that participated in the contamination and redox processes, as well as on the processes in the dyke and feeder systems that led to surface-near mineralisation with native iron and sulfides. for about 100 years (1871−1971), only two intrusions with native iron and sulfide were known: the uiffaq and illukunnguaq dykes. an extensive literature exists on the intrusions; the history of the discovery of the intrusions and their native iron bodies is reviewed below in the sections on the individual intrusions. intrusions in dyke system a hammer dal complex. this complex illustrates the composite nature of many of the native-iron-bearing intrusions, with an early intrusive pulse of contaminated basalt followed by a later pulse of much more contaminated andesite. the subhorizontal funnelor tube-shaped intrusions interconnected by thin dykes illustrate better than anywhere else the gravitational control of the mineral accumulations, with deposition of the heavy native-iron and sulfide bodies from the basaltic pulse along the lower depressions of the tubes, and accumulation of the much 50 ? 40 30 20fig. 176 10 0 m siliceous basalt native-iron cumulate xenoliths in siliceous basalt matrix sulfide-mineralised scoria tilted basalt lava flows, dip 12–15°w with columnar jointing xenolith-bearing andesite with blocky jointing with platy jointing gv03_03_065_lml s s s fig. 175. funnel-shaped intrusion in the northern segment of the hammer dal complex (fig. 174). see text for description. photogrammetrically constructed cross section, after ulff-møller (1977). for locality, see fig. 103, loc. 10. 198198 lighter sediment xenoliths below the hanging walls and roofs of the intrusions. the hammer dal complex was first discovered in 1972 in the northern wall of hammer dal about 10 km from the west coast of disko (pedersen 1975a). it has been described by ulff-møller (1975, 1977, 1983). it displays spectacular cumulates of native iron and nickeliferous pyrrhotite, and accumulations of sediment xenoliths that have floated in the magma. the complex comprises a number of intrusions, which were emplaced into a succession of basalt lava flows of the rinks dal member. the intrusions are situated along a zone that is now oriented 155–160°/80°w after a later westward tilting of the lava succession, with a horizontal extent of at least 7 km. the complex also comprises a downfaulted transgressive sill about 0.5–1 km east of the main intrusions. a block diagram of the complex is shown in ulff-møller (1977, fig. 11). the main intrusions form westerly tilted, funnelshaped, columnar-jointed masses which vary in thickness from 10 to 40 m; the individual intrusions are interconnected by thin basaltic dykes. the main intrusions widen to between 50 and 100 m in the uppermost, poorly exposed zone, which is probably close to the eruption site that may be a neck or a short feeder dyke. the best-exposed part is the lowermost intrusion (figs 175, 176). the inclined body of this has a vertical extent of 129 m and a maximum width of 30 m. it has a marginal zone up to 2–3 m wide of silicic magnesian basalt, in mf168_11cmakp_1972_01_04hdcompudmod_c.eps fe-andesite + xenoliths s fe cum fe-basalt fig. 176. the lower part of the funnel-shaped intrusion in the hammer dal complex. s: the sulfide-mineralised zone. fe cum: the nativeiron cumulate zone in fig. 175. the position of the photo is indicated in fig. 175. 199 50 40 30 20 10 0 m gv03_03_066_lml siliceous basalt native-iron cumulate xenoliths in siliceous basalt matrix sulfide-mineralised scoria tilted basalt lava flows, dip 12–15°w late dyke with columnar jointing xenolith-bearing andesite with blocky jointing with platy jointing s s mf170_17cmakp_1972_02_15hdmod2_c.eps fe s fe-andesite 11 m fe s fe-andesite 11 m d fig. 178. transgressive sill east of the main funnel-shaped intrusion of the hammer dal complex. the sulfide mineralisation (s) and the native iron cumulate zone (fe) are indicated. the cross-cutting dyke (d) in the right part of the photo is also seen in the right part of fig. 177. fig. 177. transgressive sill east of the main funnel-shaped intrusion of the hammer dal complex. see text for description. photogrammetrically constructed cross section, after ulff-møller (1977). for locality, see fig. 103, loc. 11. 200200 which native iron accumulated in local depressions. the central part consists of magnesian andesite with native iron (table 11; ulff-møller 1983, table 7). the transgressive sill intrusion east of the main intrusion is up to 10 m thick and is cut by younger, uncontaminated basalt dykes (figs 177, 178). the sill is composite, like the main intrusions, and has a lower basaltic part with native-iron accumulation and an upper basaltic part; the central part is a fairly silicic graphite andesite with native iron (ulff-møller 1983, table 7). along the lower margin of the transgressive sill, where it cuts across the vesiculated top zone of a basaltic lava flow, a less than 1 m thick zone is developed that is heavily mineralised with nickeliferous pyrrhotite. the basaltic upper margin of the sill is strongly contaminated and packed with sediment xenoliths. inspection of these sediment accumulations reveals a strong dominance of graphitic, magma-equilibrated mudstone, while sandstone is scarce (fig. 179). hanekammen complex. the hanekammen complex was discovered in 1972 (pedersen 1975a) and later investigated by f. ulff-møller (pedersen & ulff-møller 1980, mf171_8cmakp_1972_h_hdalxennatverdmod_c.epsfig. 179. sediment xenoliths packed by floating along the upper margin of the transgressive sill of the hammer dal complex (fig. 177). the xenoliths are almost all magma-equilibrated mudstone with graphite. length of pencil 13 cm. mf172_17cma1972_01_29mod_c.epsfig. 180. volcanic neck of the hanekammen complex, seen from the west. the 300–400 m wide neck is shown in the upper left part of the map area in fig. 181. 201 ulff-møller 1983, p. 28–55 and ulff-møller 1990). the complex comprises a number of intrusions. the northernmost part is a prominent volcanic neck c. 2.5 km north-west of hanekammen (figs 180, 181). it measures 300 × 400 m in cross section and probably represents an eruption site. the neck is composed of basaltic andesite with native iron and is rich in sediment xenoliths. a series of nnw–sse-trending dykes extends from this neck for at least 13 km towards south-south-east. in the northern 3 km there are several irregular dyke intrusions, whereas the southern 10 km is a simple en echelon dyke which has its southernmost exposure in the western wall of stordal. the composite intrusions have early marginal pulses ranging from native-iron-bearing, contaminated basalt to 500 m gv01_02_164_lml fig. 182 fig. 180 fig. 181. geological map of an area on north-west disko with the hanekammen complex necks and dykes (red) and the nordfjord complex neck and dykes (blue). the surrounding lava succession is the rinks dal member; pale brown: fe-ti rich basalt, yellow: aphyric and feldspar-phyric basalt, green: feldspar-phyric basalt. the red lava remnant on the peak hanekammen is a nordfjord member flow. dark grey is moraine and white with blue margins is ice. map by finn ulff-møller (1983). the map location is shown in fig. 174. 202202 almost uncontaminated basalt, whereas the later pulses consist of native-iron-bearing basaltic andesite. the hanekammen complex has been investigated in detail on the se-exposed wall of a large ridge extending from the peak hanekammen towards the south-west. here, several tube-shaped or irregular, horizontal intrusive bodies are connected into a subvertical dyke that widens upwards to form an eruption site with crater breccia on the top of the ridge (fig. 182). the lowermost intrusive body is particularly instructive (fig. 183; table 11). it is an irregular intrusion with a marginal zone of native-iron-bearing, silicic basalt which contains a small native-iron cumulus zone, and a later central pulse of native-iron-bearing basaltic andesite which contains chemimf174_11cmfum_1987_04_12mod_c.eps d d cr fig. 183 fig. 180 fig. 182. dyke and crater of the hanekammen complex viewed from the south-southeast along the strike of the near-vertical dyke. the dyke (d) is seen here to widen upwards into an eruption site with an oxidised crater breccia (cr) on the crest of the ridge in the foreground. the larger volcanic neck in fig. 180 is visible on the ridge in the background. compare with the map in fig. 181. photo: finn ulff-møller. 203 cally zoned native-iron bodies. this intrusion was used by ulff-møller (1990) to discuss the relation between zoning in the native iron and the dynamics of the process of contamination with crustal rocks. nordfjord complex. the nordfjord complex was discovered in 1979 (pedersen & ulff-møller 1980) and described by ulff-møller (1983, p. 56–65). a small outcrop on the north coast of the innermost part of nordfjord, consisting of silicic basalt with a very rich assemblage of sediment xenoliths, had already been discovered in 1968 (pedersen 1969). exposures are fragmentary. the main exposure has a n–s extent of 1 km and is found on both the northwestern and south-eastern wall of the ridge extending south-west from the hanekammen peak (fig. 181). it comprises a composite dyke up to 10 m wide that has locally developed upwards into a brecciated crater pipe with a circular outline and a diameter of c. 60 m. just 70 m below the crater, a tubular subhorizontal intrusion is exposed that has a height of 35 m and a width of at least 25 m. it is composite, and its basaltic marginal part (table 11) contains small native-iron bodies and a zone with 10–15 cm-sized, coarse-grained plagioclase-olivine nodules. the inner part of the intrusion is a younger pulse of basaltic andesite to andesite with native iron (table 11). about 2 km south of the crater, on the south wall of rink dal, the complex continues as a funnel-shaped, subhorizontal intrusion that shows spectacular columnar jointing (fig. 184). the core of this part of the intrusion is a native-iron-bearing silicic basalt (ulff-møller 1983, table 5, no. 3) similar in composition to the silicic basalt 10 m wsw ene silicic basalt native-iron cumulate basalt lavas basaltic andesite with native iron vesiculated zone gv03_03_067_lml.epsfig. 183. the lowermost intrusive body of the hanekammen complex. see text for description. the arrow points to an accumulation of sediment xenoliths in the marginal basalt. drawing modified from ulff-møller (1990). location indicated in fig. 182. mf176_8cmfum1983thesisfig20mod_c.epsfig. 184. the nordfjord complex in the south wall of rink dal, north-western disko. the complex here forms a subhorizontal, funnel-shaped intrusion with spectacular columnar jointing. the height of the intrusion tube seen is 20 m. photo: finn ulff-møller. 204204 found on the north coast of nordfjord (pedersen 1969). this indicates that the nordfjord complex extends for at least about 12 km. the chemical compositions of the rocks of the nordfjord complex are very similar to the most common magnesium-poor basalts and basaltic andesites represented by most of the contaminated lavas of the nordfjord member. stordal dyke with 10 tons native-iron boulder. for the first time in more than 100 years of research, a new major boulder of native iron was discovered in 1985 (ulff-møller 1986). the locality is in the northern wall of stordal about 3.5 km east-south-east of point 780 m at 389 m altitude (central disko section at 33.2 km; fig. 174). the boulder belongs to a 2 m thick, contaminated basic dyke from which it is slightly displaced due to sliding. it measures 0.8 × 1.5 × 2 m and has an estimated weight of about 10 tons (fig. 185). a modal and chemical comparison with the large uiffaq native-iron boulder (table 10; ulffmøller, unpublished data 1990) shows that the stordal boulder contains much more troilite (fes) and much less cohenite (fe3c) than the uiffaq boulder. the host dyke trends nw–se and forms part of a small swarm of two to three thin dykes that cut lavas and hyaloclastites of the vaigat formation but not the underlying disko gneiss ridge. the dyke belongs to dyke mf177_17cml922721_c.epsfig. 185. the 10 tons native iron boulder found in stordal in 1985 (fig. 174). the boulder measures 0.8 × 1.5 × 2 m and comes from a 2 m thick, contaminated basic dyke from which it is slightly displaced due to sliding. the substrate in the photo is a lava flow. dyke iron block lava flow lava flow gv03_03_069_lml.epsfig. 186. a model for native iron accumulation in steep intrusions from droplet size to boulder size by gravity accumulation (arrows) at irregularities such as on ledges and in wedges. after ulff-møller (1986). 205 system a and is interpreted as the result of horizontal magma transport from a high-level magma reservoir situated within the sedimentary basin north-west of the disko gneiss ridge. the dyke rock is olivine-phyric and contains native iron and many xenoliths of graphite-rich, magmamodified mudstone. its composition varies between silicic basalt and basaltic andesite (table 11, samples f062, f066). a model for native-iron accumulation from drops to block size is shown in fig. 186 (ulff-møller 1986). inner giesecke dal dyke zone. this is the easternmost part of dyke system a. the zone is rather narrow and extends from east of giesecke dal in the north-west for more than 22 km to stordal in the south-east, where it lies east of the stordal dyke described above (fig. 174). the zone hosts a number of nw–se-trending, contaminated dykes that are basaltic at the north-western and south-eastern ends and composite with native-iron-bearing andesitic cores within an 8 km long central part of the zone. the central part is therefore interpreted as the primary upwelling zone for the strongly contaminated magmas from a reservoir within a succession of mudstones and sandstones. these dykes have been investigated and sampled by f. ulff-møller (ulff-møller 1991), but the results have not been published before. the northernmost dyke in this zone has been sampled 2–3 km north-west of point 1460 m north of giesecke dal (fig. 103, loc. 28). it is a nw-trending, 4–5 m thick, regularly columnar-jointed dyke of contaminated basalt (fig. 187) with about 8.9 wt% mgo. the southernmost dyke in the zone cuts the disko gneiss ridge and the vaigat formation (fig. 174). it is 8–10 m thick and has fine-grained margins and a crumbling core of more coarse-grained dolerite; from a distance, this gives it the appearance of a double dyke. it consists of contaminated basalt with 11.3 wt% mgo and is rich in graphite-bearing magma-modified mudstone xenoliths 2–7 cm in size. the central part of the dyke zone hosts one major dyke and several smaller ones, which are all composite with native iron. within 1 km of the main dyke, the smaller dykes may cut each other at low angles. locality stordal uiffaq year of find 1985 1870 weight, tons 10 25 estimated mode, vol.% iron 60 43 cohenite 7 54 troilite 30–40 3 schreibersite 0 trace silicate glass 0.1 0 chromite 0.01 0 wüstite trace ? calculated chemical composition, wt% fe 80 91.6 co 0.4 0.5 ni 2 1.8 cu 0.1 0.16 c 0.5 3.62 p 0.4 0.15 s 10–15 1.09 o trace 0.97 source: finn ulff-møller, personal communication, c. 1990. table 10. estimated compositions of native iron boulders from disko mf179_8cmfum_1985_12_10mod_c.eps fig. 187. the northernmost contaminated dyke in the inner giesecke dal dyke zone, 2–3 km north-west of point 1460 m north of giesecke dal (fig. 174). length of hammer handle 60 cm. photo: finn ulff-møller. 206206 inner giesecke dal intrusion. the single major dyke in the inner giesecke dal dyke zone is named the inner giesecke dal intrusion. the northernmost, well-exposed part of the inner giesecke dal intrusion is located c. 500– 600 m west of point 1560 m north-east of giesecke dal (fig. 103, loc. 29), where it forms a spectacular, nativeiron-bearing composite dyke about 10 m thick. the dyke is rich in sediment xenoliths and its core is strongly contaminated and transitional between basaltic andesite and andesite in composition. in the same area a loose boulder with a native-iron cumulate zone more than 0.5 m thick was located, together with other loose blocks with kidney-shaped native-iron bodies with troilite rims 2–3 cm in size. the central part of the inner giesecke dal intrusion is exposed around the top of the high plateau at point 1590 m between giesecke dal and morten porsild dal (fig. 103). here, the intrusion is 12–15 m wide and dips c. 70° ne; it has a distinctly rusty colour due to weathering of native iron and sulfide. the dyke is composite with 1–2 m thick basaltic margins and an andesitic core; as an unusual feature, it also contains an inner zone of basalt 1 m thick, which is slightly chilled against the andesite. the marginal basalt contains small native-iron cumulates along the footwall and sediment accumulations along the hanging wall. another, parallel, composite dyke with native iron occurs c. 10 m south-west of the main dyke. the southernmost exposures of the inner giesecke dal intrusion are situated c. 2 km north-west of point 1650 m north of agatfjeldet, on walls at the end of morten porsild dal (fig. 174). here the intrusion forms a c. 7 m thick, segmented, composite dyke dipping 70° ne. it has a basaltic to basaltic andesitic margin c. 0.5 m thick, and an andesitic core with up to 1 cm large native-iron bodies (table 11, nos f045 and f048; fig. 188). on both sides of the composite intrusion, there are several 1–4 m thick, subparallel basaltic intrusions with sediment xenoliths, and also other intrusions that are only slightly contaminated. mf180_17cmfum_1985_05_31mod_c.eps picrite flows picrite flows fig. 188. the southernmost exposures of the inner giesecke dal intrusion, c. 2 km north-west of point 1650 m north of agatfjeldet, at the end of morten porsild dal (fig. 174). the intrusion forms a segmented composite dyke c. 7 m thick, with basaltic to basaltic andesitic margins and an andesitic native-iron-bearing core. photo: finn ulff-møller. 207 outer giesecke dal dyke. a prominent composite dyke intrusion was discovered by f. ulff-møller in 1984 c. 8 km south-east of the entrance to giesecke dal and c. 1.5 km east of the volcanic neck at point 882 m (fig. 103, loc. 30). the dyke trends wnw–ese and is well exposed on both walls of a gully south of giesecke dal from where it can be traced eastwards for 3 km. the dyke is up to 25 m thick and composite, with a marginal zone of contaminated basalt up to 7 m thick, which locally contains a native-iron cumulus zone several metres thick, and a core zone of basaltic andesite with native iron and sediment xenoliths. small swarm of nw–se-trending dykes around innermost hammer dal. a small swarm of at least five nw– se-trending dykes discovered by f. ulff-møller in 1985 and 1987 in the innermost hammer dal area belongs to dyke system a. a ‘double dyke’ (two closely set parallel dykes) has a main exposure c. 3 km nnw of point 1380 m (fig. 103, loc. 31). the western dyke is 6–8 m thick and distinctly contaminated (table 11, no. f074). it is notable for a rich assemblage of xenoliths up to 1 m large of magmamodified sandstone and mudstone and abundant, 2–50 cm large gabbroic nodules. the eastern dyke is 20–30 m thick and only slightly contaminated (table 11, no. f076). it is composed of mediumto coarse-grained dolerite with light grey residual veins and up to 5 cm large gabbroic xenoliths. about 1.5–2 km east of the double dyke is another parallel basaltic dyke (fig. 103, loc. 32; table 11, no. f162) which is c. 5 m thick and contains glomerocrysts up to 1 cm large of plagioclase and abundant xenoliths of sandstone and mudstone. along the strike and across a glacier, 8 km to the south-east, loose blocks of contaminated basalt from a dyke occur in the talus close to steenstrup dal. the basalt contains coarse plagioclase and olivine glomerocrysts and abundant small xenoliths of sandstone and mudstone and is similar to, but less evolved than, the dyke 8 km to the north-west. about 1.5–2 km farther to the east another composite dyke c. 4 m thick is exposed a few kilometres south-west of point 1610 m (fig. 103, loc. 33). it strikes nnw–sse and is composed of a distinctly flow-laminated, contaminated basalt with a core of more silicic, contaminated basalt with disseminated graphite (table 11, no. f169). a similar, flow-laminated dyke has been traced a few kilometres to the north-west and is probably the same dyke. the flow-lamination is very irregular and discordant to the strike direction of the dyke. finally, a nw–se-trending dyke has been traced south-east of the point 1610 m plateau between the innermost parts of hammer dal and giesecke dal (fig. 103, loc. 34). this dyke is a c. 10 m thick, contaminated, evolved dolerite (table 11, no. f173) with scattered, c. 1 cm large xenoliths of plagioclase and graphite, i.e. magma-modified mudstone. intrusions in dyke system b killiit dyke. the third discovery of a native-iron-bearing intrusion came almost 100 years after the uiffaq report, when fundal (1972, 1975) described an 8 m thick basalt dyke with native iron and plagioclase-spinel-graphite xenoliths cutting the disko gneiss ridge at killiit (old spelling kitdlît) on the south coast of disko, 9 km west of qeqertarsuaq town (figs 27, 36). another segment of the dyke at luciefjeld c. 7 km north of the coast cuts basaltic lavas of the rinks dal member but not the underlying gneiss. based on these observations fundal (1975) argued that the native iron and the graphite-bearing xenoliths could not be a product of sediment–magma reaction but required a much deeper origin. the same arguments were subsequently pursued by bird and weathers (1977) but were later abandoned by bird et al. (1981). the killiit dyke was subsequently mapped northwards along the disko gneiss ridge over a distance of more than 80 km (pedersen 1977c; ulff-møller 1979; pedersen & ulff-møller 1980; ulff-møller 1983, 1985). in the south the dyke is basaltic (e.g. pedersen 1979b table 2) and strikes close to n–s. about 65 km north of killiit it changes direction to nw–se and becomes composite with a basaltic margin and an andesitic core (table 11; ulff-møller 1985, table 1, nos 2 and 3; fig. 189). the dyke increases in thickness to 10–15 m towards the north and was clearly emplaced from the north-west. in the northern area, there are several offshoots from the dyke that consist of basalt packed with gneiss xenoliths (fig. 190). in a limited area, the dyke develops an oxidised, crater-like breccia along one of its margins. from just north of kangerluk, a c. 100 kg native-iron body called the kitdlît lens has been described in detail and reveals a complex solidification history (ulff-møller 1985). this paper is a major contribution to our knowledge about the native iron. the present interpretation of the killiit dyke is that it originated from a high-level magma reservoir within the sedimentary basin north-west of stordal as a part of dyke system a. the intrusion was emplaced in several pulses, of which the first was a native-iron-bearing basalt magma 208208 that travelled subhorizontally for more than 85 km towards the south to killiit, where basalt with sediment xenoliths and native iron has now been found cutting the disko gneiss ridge. when the dyke changed direction to n–s, about 65 km north of the coast it became part of dyke system b. shortly afterwards, a second pulse of much more contaminated and viscous andesite magma with native iron was emplaced towards the south but became exhausted after having travelled only 15–20 km. intrusions in dyke system c maniillat dyke. the maniillat dyke trends nw–se and runs subparallel with the vaigat coast on north-eastern disko for at least 12 km between maniillat and pyramiden (fig. 174). the dyke dips 70–80° ne and is 5–6 m thick in its northern part and around 10 m thick farther south. it cuts both picritic lavas of the vaigat formation and basalt flows of the rinks dal member of the maligât formation. the dyke has marginal zones of contaminated basalt (table 11, no. 138235) with olivines up to 2 cm in size and native-iron bodies up to 1 cm in mf181_8cmfum_1979_11_17mod_c.eps fe fe fig. 189. the northern part of the killiit dyke, 65 km north of killiit. the dyke (fe) is here c. 15 m thick and cuts lava flows of the rinks dal member. it is composite with a basaltic margin and a c. 13 m thick andesitic core; both parts are native-iron-bearing. locality c. 10 km south of stordal and c. 1 km west of point 1266 m (figs 6, 174). photo: finn ulff-møller. fig. 190. offshoot from the killiit dyke 75 km north of killiit. the dyke here trends nw and cuts lava flows of the rinks dal member. the apophysis is c. 1 m thick and consists of basalt packed with xenoliths of gneiss or arkose. locality c. 1 km east of point 1578 m in the north-western side of a corrie leading into the south side of stordal (figs 6, 174). photo: finn ulff-møller. mf182_8cmfum_1979_05_15mod_c.eps 209 size, beside a range of sediment xenoliths and also sharpedged xenoliths of uncontaminated basalt with plagioclase phenocrysts. the central, 4–6 m wide part of the dyke (table 11, no. 138233) has a sharp, but not glassy, contact against the marginal basalt and is composed of andesite with microphenocrysts of olivine and orthopyroxene; it also contains disseminated graphite and xenoliths of magma-modified mudstone. along the hanging wall, the marginal basalt is particularly rich in sediment xenoliths, and at one locality it contains a 40 cm thick zone of densely packed xenoliths of sandstone, mudstone and coal fragments with an appearance almost like a sediment breccia (fig. 191). the dyke disappears beneath the glacierand moraine-filled areas around qullissat. it is possibly connected to the small dyke segment at illukunnguaq c. 35 km to the south-east (see below). illukunnguaq dyke. the second early known intrusion is a 5 m thick basic dyke intruding sandstone of the atane formation near the abandoned settlement illukunnguaq (igdlukúnguaq) on eastern disko. the dyke strikes nw–se and can be followed for about 800 m close to the vaigat coast (steenstrup 1874, p. 88, also briefly mentioned by nordenskiöld 1871). the dyke contained a 28 tons body of nickeliferous pyrrhotite (now nearly mined out) that was described in detail by pauly (1958). the dyke also contains a range of sediment xenoliths (fundal 1975). the occurrence shows that the contaminated intrusions have a potential for sulfide mineralisation. the dyke has been described by ulff-møller (1983, p. 92–98) who found sulfide blebs with native iron in its glassy chill zone. an analysis is presented in table 11, no. 362140. the dyke is now interpreted as part of dyke system c, and it is possible that it acted as feeder for the composite native-iron-bearing lava flow in the lower niaqussat member on eastern disko (fig. 10, profile 8; fig. 140). other dyke intrusions uiffaq dyke. the discovery of three large native-iron boulders (25, 6.5 and 4 tons) on the south-western shore of the uiffaq peninsula west of qeqertarsuaq was reported by nordenskiöld (1871) who interpreted them as iron meteorites. their unusual composition was reported by nordström (1871). the history of recovery and research of the native iron is summarised by sjögren (1916) and an extensive bibliography is given by bøggild (1953). a memoir based on the largest (25 tons) iron boulder housed in stockholm (löfquist & benedicks 1941) showed that the native iron is dominated by iron carbide (cohenite) and is a cast-iron type. field observations by nauckhoff (1872, p. 38) and steenstrup (1875, fig. 3) suggested the presence of a poorly exposed basaltic dyke at the shore. the host body for the native-iron boulders was finally proved by magnetometer mapping to be a c. 7 m thick, approximately n–s-running, composite basaltic dyke (fundal 1975). various segments of the dyke have been mapped over a distance of 8 km across the uiffaq peninsula to its north coast (pedersen et al. 2000); the dyke is shown in fig. 81. detailed data on the uiffaq native iron and sulfide have been presented by goodrich (1984), goodrich & bird (1985) and howarth et al. (2017). mf183_11cmakp_1971_06_18mod_c.eps fig. 191. the maniillat dyke (fig. 174) with an accumulation of closely packed, black sediment xenoliths along its hanging wall. length of ruler 1 m. 210210 ‘rink’s dyke’. for many years, the existence of this dyke was inferred from a sample collected by h. rink around 1850 on the north coast of kangerluk. the dyke was rediscovered by f. ulff-møller in 1978 (fig. 174) and briefly described by him (ulff-møller 1979, 1983). it is only known to be exposed at the beach where it is vertical, c. 4 m thick and strikes n–s. it is composite, with a marginal zone of silicic basalt, a hybrid zone 10–20 cm thick, and a central part 2.5 m thick of basaltic andesite (table 11). the central part contains numerous 1–10 cm equilibrated sediment xenoliths consisting mainly of plagioclase and graphite. no native iron was observed. mellemfjord dyke. this dyke was discovered by akp in 1974 at saqqarliit ilorliit on the north coast of mellemfjord (fig. 174) and described by ulff-møller (1979, 1983). the dyke is exposed over about 1 km where it is vertical and strikes n–s; it is c. 10 m thick at low altitudes but narrows to just 1 m at 700 m a.s.l. the dyke is rich in vesicles. chemically it is a basaltic andesite (table 11), probably with a more basic margin. the chilled margin contains sulfide blebs but no native iron. concluding remarks on the strongly contaminated dyke systems the largest and most frequent strongly contaminated intrusions are situated on western and north-western disko west of the disko gneiss ridge. therefore, the high-level magma reservoirs within the underlying sediments must have been concentrated in this area, and possibly also beneath the shallow sea north-west of disko. in addition to the described contaminated intrusions, there are other minor exposures on disko west of the disko gneiss ridge, and a detailed search, particularly on south-western disko, would doubtlessly increase the known localities. however, the general picture will most probably not change. chemical compositions of the nordfjord and niaqussat member dykes and feeder systems the intrusions range from only slightly contaminated silicic basalts to strongly crustally contaminated andesites; there are no dacites or rhyolites. the chemical compositions are illustrated in fig. 192, and representative analyses are shown in table 11. most of the contaminated intrusives have compositionally similar counterparts in the lavas of the nordfjord and niaqussat members (figs 145, 147, 167, 169), but some have not, most notably the killiit dyke which has higher mgo for a given sio2 than any other intrusions or lavas. the weakly contaminated dykes on eastern disko and nuussuaq (green dots in fig. 192) represent an almost normal magma evolution series. in contrast, the large intrusive dyke and crater complexes, such as the nordfjord complex, the hanekammen complex and the killiit dyke, show different compositional evolutions, mainly caused by different compositions of the starting magmas of each system. the killiit dyke had the most magnesian starting magma (12–13 wt% mgo), the hanekammen complex was intermediate (9–10 wt% mgo) and the nordfjord complex was the most evolved (c. 8 wt% mgo). the contamination processes comprised sidewall melting, mixing, mineral re-equilibration and reduction (e.g. ulff-møller 1990; larsen & pedersen 2009), and in each system the contamination led to increasing sio2 and k2o and decreasing mgo, cao and feo*, while tio2 and p2o5 stayed approximately constant. the cr contents are higher than in uncontaminated magmas because chromite fractionation was delayed by the contamination process (pedersen 1985), but all the intrusions have a lower degree of cr-enrichment than the most contaminated rocks of the vaigat formation (the asuk and kûgánguaq members). also, no intrusions show ni-depletion as severe as in the asuk and kûgánguaq members, but some native-iron-bearing samples of the hanekammen complex and the stordal dyke show ni-accumulation, presumably in the metallic phase. there are no compositional differences between dyke systems a, b and c described above. dyke groups such as the dykes in giesecke dal and inner hammer dal show widely variable compositions, and the individual dykes must have been generated independently of each other. there are not enough chemical data to characterise the hammer dal complex further. 211 45 47 49 51 53 55 57 59 61 63 65 0 100 200 300 400 500 600 700 800 900 1000 1100 cr (ppm) 6 8 10 12 14 16 feo* 1 2 3 4 weakly contaminated, outside w disko various contaminated, c-w disko uiffaq dyke killiit dyke stordal fe-dyke nordfjord complex hanekammen complex hammer dal complex inner hammer dal dykes giesecke dal dykes maniillat dyke illukunnguaq dyke tio2 10 100 1000 ni (ppm) 5 7 9 11 13 15 cao kûgánguaq mb kûgánguaq mb asuk mb asuk mb rdm + vf rdm + vf asuk mb sio2 0 2 4 6 8 10 12 14 0 2 4 6 8 10 12 14 0 2 4 6 8 10 12 14 0 2 4 6 8 10 12 14 0 2 4 6 8 10 12 14 0 2 4 6 8 10 12 14 mgo (wt%) mgo (wt%) fig. 192. variation diagrams of selected major and trace elements for contaminated intrusions of the nordfjord and niaqussat members. major element data in wt% oxides recalculated volatile-free. feo* is total iron as feo. c–w: central to west. arrows indicate compositional changes with increasing contamination in the killiit dyke and the nordfjord and hanekammen complexes. the fields for the contaminated asuk and kûgánguaq members of the vaigat formation are shown for comparison in the cr and ni diagrams. the shaded areas in these two diagrams labelled rdm + vf are the fields for the uncontaminated rinks dal member and the uncontaminated parts of the vaigat formation. 212212 6261 6261 6161 6161 6271 6271 6281 6281 6291 6291 6281 6281 175950 176007 264355 264424 264350 264349 f062 f066 f045 f048 f074 f076 5434.27 5434.27 5419.78 5419.86 5421.8 5421.8 5355.62 5355.62 5404.30 5404.30 5423.43 5423.39 7009.14 7009.15 7002.99 7003.00 7002.7 7002.7 6959.75 6959.75 7006.40 7006.40 7006.91 7006.91 472 490 1055 1100 1015 1030 385 386 1223 1222 592 587 49.15 57.55 50.38 53.92 48.42 56.10 50.09 53.01 52.10 58.61 51.00 47.92 1.37 1.30 1.47 1.43 2.97 2.10 1.281 1.237 1.761 1.281 2.095 1.513 13.49 14.35 14.17 16.30 13.54 15.70 14.33 13.91 14.66 15.57 14.93 15.00 0.88 0.00 1.22 0.00 4.06 2.55 11.79 11.18 10.74 8.22 12.11 11.65 10.52 8.40 9.96 8.73 9.88 7.14 0.16 0.12 0.21 0.13 0.145 0.135 0.154 0.116 0.166 0.176 10.62 5.85 9.37 6.22 5.92 3.38 9.10 8.86 6.84 5.59 6.19 7.94 10.00 6.97 9.84 7.98 10.91 7.02 8.94 7.28 9.35 6.41 9.52 12.55 1.66 2.17 1.99 1.84 2.54 2.32 1.73 1.73 2.05 2.28 2.19 2.04 0.310 1.400 0.330 0.900 0.260 1.250 0.604 0.845 0.756 1.021 0.491 0.177 0.120 0.150 0.150 0.180 0.310 0.280 0.156 0.170 0.201 0.187 0.217 0.132 1.63 1.75 1.06 2.09 1.03 1.95 1.53 1.14 0.86 0.35 0.63 0.54 99.75 99.52 100.10 99.71 100.05 99.92 99.69 99.49 99.47 99.64 99.55 99.64 11.31 8.40 11.06 8.73 13.53 9.43 10.61 10.06 9.66 7.40 10.90 10.48 65.50 58.48 63.15 59.04 46.94 42.02 63.44 64.06 58.87 60.43 53.47 60.51 79.1 55.7 72.0 53.0 181 99.2 65.4 76.0 87.8 64.2 94.3 73.4 138 119 148 127 294 114 59.8 171 49.4 55.9 105 161 548 161 343 228 78.9 115 151 482 128 84.0 84.5 152 33.8 27.4 31.0 26.9 37.1 30.0 33.5 28.7 32.6 25.7 34.7 37.3 284 187 258 213 432 209 308 244 266 185 279 323 859 450 665 429 120 160 936 742 601 388 303 386 18.7 19.5 19.6 21.8 24.1 24.9 16.8 19.3 20.9 21.2 22.3 19.4 6.76 32.2 7.54 26.3 4.77 33.8 3.48 15.8 13.5 39.2 23.2 2.61 165 185 191 230 243 238 181 192 199 218 215 199 23.5 27.5 23.2 24.2 39.8 34.2 16.5 21.3 24.8 25.3 30.9 23.6 96.6 192 111 155 181 218 64.5 114 132 190 155 81.1 4.14 7.84 4.90 7.27 9.08 10.25 2.71 5.13 6.02 8.37 7.32 4.16 0.237 0.805 0.220 0.763 0.135 0.814 0.128 0.344 0.327 1.04 0.434 0.038 62.01 304 127 270 53.79 348 39.4 148 145 340 163 28.8 7.02 19.5 10.4 18.0 10.1 23.4 4.68 11.7 12.8 21.9 13.3 4.56 17.16 42.4 24.1 39.6 28.4 51.7 11.9 26.8 30.1 48.3 31.8 12.8 2.48 5.37 3.19 4.80 4.46 6.60 1.67 3.47 3.85 5.73 4.29 2.06 11.8 21.7 14.3 19.8 22.3 27.9 8.24 15.2 17.0 23.1 19.5 10.7 3.35 4.95 3.72 4.38 6.41 6.37 2.34 3.72 4.20 4.96 5.05 3.43 1.11 1.19 1.19 1.30 2.20 1.65 0.94 1.18 1.30 1.30 1.60 1.25 4.05 5.35 4.38 4.64 7.37 6.75 2.84 4.02 4.60 5.11 5.74 4.16 0.674 0.830 0.694 0.741 1.20 1.05 0.471 0.652 0.750 0.798 0.930 0.688 4.07 4.90 4.32 4.24 7.38 6.24 2.90 3.91 4.48 4.47 5.51 4.31 0.832 0.989 0.847 0.853 1.46 1.23 0.574 0.775 0.906 0.917 1.12 0.900 2.26 2.64 2.36 2.36 3.91 3.32 1.57 2.07 2.43 2.47 3.01 2.36 0.323 0.393 0.329 0.334 0.551 0.482 0.216 0.306 0.357 0.364 0.426 0.354 1.98 2.34 1.99 2.06 3.30 2.95 1.35 1.82 2.09 2.16 2.63 2.07 0.309 0.370 0.297 0.310 0.510 0.445 0.196 0.264 0.311 0.326 0.391 0.313 2.52 4.84 2.94 3.98 4.59 5.61 1.74 3.10 3.51 4.92 4.05 2.14 0.394 0.516 0.453 0.656 0.852 0.898 0.210 0.338 0.405 0.569 0.481 0.272 2.29 5.23 2.39 4.00 6.54 8.66 1.44 5.35 3.98 5.89 4.28 0.610 1.15 4.59 2.10 4.11 0.799 5.03 0.735 2.42 2.64 5.15 2.52 0.345 0.320 1.18 0.509 1.00 0.243 1.23 0.189 0.603 0.680 1.25 0.650 0.113 margin centre margin centre margin centre margin centre lith. code ggu no. deg. w deg. n altitude, m sio2 tio2 al2o3 fe2o3 feo mno mgo cao na2o k2o p2o5 volatiles sum feo* mg number zn cu ni sc v cr ga rb sr y zr nb cs ba la ce pr nd sm eu gd tb dy ho er tm yb lu hf ta pb th u for explanation of lithological codes, see table 1. for petrographical notes on the samples, see table 11c. geographical coordinates in wgs 84. first two digits are degrees, then follow minutes in decimal form. major elements in wt% (xrf analyses). trace elements in ppm (zn–ga: xrf analyses; rb–u: icp-ms analyses). feo* = total iron as feo. mg number = 100 × atomic mg/(mg+fe2+), with the iron oxidation ratio adjusted to fe2o3/feo = 0.15. table 11a. chemical analyses of intrusions of the nordfjord and niaqussat members intrusion locality hammer dal complex hanekammen complex nordfjord complex inner giesecke dal inner hammer dal swarm stordal dyke hammer dal hanekammen rink dal giesecke dal inner hammer dalstordal 213 6281 6281 6281 6171 6171 6281 6281 6281 6281 6281 6281 6281 f162 f169 f173 176669 264338 138235 138233 362140 176683 175156 175164 175005 5420.36 5417.18 5413.78 5346.26 5353.5 5332.25 5332.26 5234.72 5412.18 5416.2 5416.2 5434 7007.09 7006.40 7006.64 6919.306 6956.36 7010.930 7010.93 6953.103 6921.676 6934.1 6934.1 6946 778 1019 1530 514 755 1059 1066 82 1 4 4 170 49.05 51.68 50.30 52.09 56.61 48.67 56.66 51.16 50.67 50.92 53.14 54.90 2.473 1.162 2.454 1.35 1.25 1.75 1.28 1.57 1.73 1.62 1.58 1.45 13.22 14.49 12.63 13.81 14.05 13.31 15.36 13.49 14.53 14.36 15.33 14.86 14.36 10.44 13.11 0.00 1.88 3.03 0.90 1.48 3.37 1.38 1.60 0.60 9.44 5.35 8.45 6.54 8.33 7.74 8.31 7.76 8.26 0.204 0.138 0.186 0.16 0.14 0.18 0.11 0.15 0.16 0.22 0.18 0.19 5.86 9.52 7.46 10.48 8.80 10.36 6.54 9.57 7.86 8.83 6.90 7.48 10.58 8.67 10.11 9.06 7.60 10.38 6.81 9.07 10.43 10.97 8.39 8.47 2.26 1.79 2.19 1.82 2.25 2.02 2.34 1.99 2.15 2.19 2.29 2.24 0.320 0.554 0.487 0.660 0.730 0.250 0.820 0.215 0.435 0.230 0.530 0.520 0.233 0.136 0.245 0.160 0.140 0.170 0.170 0.183 0.173 0.150 0.180 0.200 1.00 1.34 0.38 0.42 0.99 1.58 2.35 2.97 0.76 1.06 2.08 0.95 99.56 99.93 99.55 99.45 99.79 99.15 98.88 100.18 100.00 100.24 99.96 100.12 12.92 9.40 11.80 9.44 7.04 11.18 7.35 9.66 10.77 9.55 9.20 8.80 47.84 67.20 56.12 69.19 71.65 65.22 64.28 66.70 59.60 65.15 60.27 63.23 110 19.3 101 76.6 49.4 95.6 77.1 88.8 84.9 82.7 75.1 77.2 174 241 94.0 36.2 70.3 162 36.1 24.3 71.8 79.1 110 39.6 52.1 733 119 42.8 102 201 76.5 123 100 143 218 104 35.2 25.9 35.7 32.2 26.4 34.5 25.3 31.2 34.6 35.0 29.7 27.0 363 221 320 259 204 327 196 265 295 296 227 231 197 823 373 854 720 492 420 613 557 622 472 586 22.5 18.7 22.6 18.4 18.0 19.3 20.9 19.7 20.2 19.7 21.2 20.5 3.89 17.2 13.0 16.1 27.6 3.34 31.6 4.29 8.97 5.37 23.0 23.1 213 205 198 194 195 188 213 240 207 196 213 236 35.4 20.9 38.2 21.9 24.3 28.2 25.1 27.0 25.8 24.7 25.0 20.1 149 119 168 123 161 116 185 137 119 105 147 135 6.74 5.07 7.29 5.30 6.56 5.30 8.11 6.24 5.08 4.40 6.73 6.31 0.054 0.398 0.218 0.448 0.659 0.107 0.839 1.08 0.160 0.126 0.542 0.714 65.8 175 94.2 141 247 47.98 300 250 97.8 60.9 205 263 8.70 12.6 10.4 10.5 16.2 7.31 20.3 11.2 8.35 7.43 15.7 13.6 23.6 28.3 27.3 24.6 35.8 19.1 44.1 27.0 20.4 18.6 35.1 29.9 3.67 3.64 4.15 3.29 4.47 2.95 5.52 3.70 2.94 2.67 4.49 3.95 18.6 15.6 20.7 14.8 18.8 14.6 22.8 16.7 14.2 12.8 19.3 17.0 5.55 3.71 6.02 3.81 4.41 4.19 4.87 4.35 3.92 3.62 4.47 3.90 1.89 1.20 1.89 1.20 1.15 1.46 1.32 1.31 1.36 1.22 1.29 1.30 6.57 4.04 7.07 4.29 4.63 4.79 5.18 4.79 4.61 4.31 4.92 4.25 1.07 0.638 1.16 0.672 0.737 0.856 0.801 0.802 0.765 0.698 0.762 0.655 6.66 3.80 7.05 3.97 4.29 4.79 4.22 4.80 4.71 4.33 4.52 3.85 1.33 0.770 1.42 0.814 0.850 0.990 0.869 0.956 0.909 0.862 0.908 0.738 3.55 1.98 3.75 2.21 2.30 2.59 2.33 2.47 2.53 2.38 2.42 1.95 0.495 0.294 0.531 0.320 0.324 0.387 0.337 0.358 0.362 0.338 0.345 0.285 2.93 1.79 3.16 1.87 2.06 2.33 2.15 2.30 2.15 2.02 2.13 1.74 0.435 0.259 0.473 0.255 0.314 0.361 0.325 0.345 0.335 0.315 0.324 0.261 4.05 3.13 4.40 3.23 4.10 3.02 4.56 3.47 3.15 2.76 3.84 3.52 0.460 0.347 0.474 1.19 0.712 0.478 0.679 0.407 0.398 0.263 0.415 0.372 1.66 1.04 2.19 2.95 2.70 1.75 6.77 3.03 2.10 1.82 5.35 4.36 0.900 2.59 1.33 2.05 3.73 0.839 4.42 2.12 1.28 1.19 3.55 2.81 0.274 0.620 0.374 0.529 0.912 0.249 1.20 0.566 0.345 0.314 0.846 0.688 margin centre margin centre margin centre lith. code ggu no. deg. w deg. n altitude, m sio2 tio2 al2o3 fe2o3 feo mno mgo cao na2o k2o p2o5 volatiles sum feo* mg number zn cu ni sc v cr ga rb sr y zr nb cs ba la ce pr nd sm eu gd tb dy ho er tm yb lu hf ta pb th u table 11b. chemical analyses of intrusions of the nordfjord and niaqussat members intrusion locality inner hammer dal swarm killiit dyke maniillat dyke rink’s dyke mellemfjord dyke uiffaq dyke illukunnguaq dyke inner hammer dal luciefjeld stordal maniilat disko fjord mellemfjordillukunnguaq uiffaq 214214 175950 olivine microphyric basalt with native iron beneath native iron cumulate. hammer dal complex, west disko. ulff-møller (1983) table 7 no. 1. 176007 andesite with native iron from the centre of intrusive body (0.37 wt% carbon). hammer dal complex. ulff-møller (1983) table 7 no. 6. 264355 plagioclase-glomerophyric basalt, 20 cm from western contact in the lowermost intrusion of the hanekammen complex. northern wall of rink dal, west disko. ulff-møller (1983) table 2 no. 3. 264424 basaltic andesite with native iron from above the centre of the lowermost intrusion in the hanekammen complex. northern wall of rink dal, west disko. ulff-møller (1983) table 2 no. 11 and ulff-møller (1990) table 1 no. 2. 264350 basalt with native iron from the margin of the northern tube-like intrusion in the nordfjord complex. northern wall of rink dal, west disko. ulff-møller (1983) table 5 no. 2. 264349 andesite with native iron from the centre of the northern tube-like intrusion of the nordfjord complex. northern wall of rink dal, west disko. ulff-møller (1983) table 5 no. 4. f062 basalt with native iron from dyke associated with 10 tons native iron boulder in the northern wall of stordal, central disko. f066 basaltic andesite with native iron from dyke associated with 10 tons native iron boulder in the northern wall of stordal, central disko. f045 basaltic andesite with native iron from margin of composite dyke (inner giesecke dal dyke), 5 km north-west of agatfjeldet, north disko. f048 andesite with native iron from centre of composite dyke (inner giesecke dal dyke), same locality as f045. f074 basalt from 6–8 m thick, nw-trending basalt dyke from the south-west side of the inner part of hammer dal, west disko. f076 basalt 2 m from the margin of 20–30 m thick, nw-trending basalt dyke from the south-west side of the inner part of hammer dal, west disko. f162 evolved basalt 0.5 m from margin of 5 m thick, nw-trending composite dyke from gully in the north-eastern wall of hammer dal, west disko. f169 basaltic andesite with graphite and sparse native iron from centre of 4 m thick nw-trending, composite dyke south-east of hammer dal, west disko. f173 basalt from near the margin of 10 m thick, nw-trending dyke between hammer dal and giesecke dal, west disko. 176669 basaltic andesite with native iron and troilite (0.09 wt% carbon). chilled margin of the killiit dyke at luciefjeld, south disko. pedersen (1979b) table 1 no. 1. 264338 andesite with native iron from the centre of the killiit dyke south of stordal, central disko. ulff-møller (1985) table 1 no. 3. 138235 basalt with native iron from margin of composite dyke (maniillat dyke), north-east disko. 138233 andesite with native iron from centre of composite dyke (maniillat dyke), same locality as 138235. 362140 magnesian basaltic andesite from nw-trending basalt dyke associated with a 28 tons pyrrhotite body with traces of native iron. illukunnguaq, east disko. 176683 basalt with native iron from the uiffaq dyke. uiffaq, south disko. 175156 basalt with native iron from composite dyke (rink’s dyke). north coast of disko fjord. ulff-møller (1983) table 1 no. 6. 175164 basaltic andesite with native iron from core of composite dyke (rink’s dyke). same locality as sample 176156. ulff-møller (1983) table 1 no. 7. 175005 basaltic andesite with native iron from dyke at saqqarliit ilorliit. mellemfjord, west disko. ulff-møller (1983) table 1 no. 4. table 11c. notes on analysed samples of intrusions of the nordfjord and niaqussat members 215 volume relations of the maligât formation in estimating the volume of the maligât formation, it must be remembered that it is not well delimited to the west and south; the formation extends below sea level west of disko and must also have extended south of disko where it is now removed by erosion. therefore, measured volumes of the formation are minimum values. on the other hand, the formation is reasonably well delimited to the north and east. the volume calculations presented here are based on a common area comprising all of disko and nuussuaq east of the itilli fault, delimited to the north by the north coast of nuussuaq, to the west by the itilli fault and longitude 55°w, to the east by longitude 52°w and to the south by latitude 69°15′n which passes through qeqertarsuaq town (see e.g. fig. 4). the area measures close to 18 000 km2. estimated volumes of the various units of the maligât formation within the delimited area are presented in table 12. the estimates are based on thickness measurements with interpolations in the five geological sections on a scale of 1:20 000 together with thicknesses measured in the many sampled profiles. the numbers were plotted on maps and the thicknesses contoured. on western disko, the succession dips below sea level and the lower units are extrapolated. in western and central nuussuaq, extrapolation is also required due to erosion. the values given in table 12 are rounded figures because of the uncertainties involved, in particular for the more voluminous units. the total volume of the formation within the delimited area is close to 21 200 km3, of which the uncontaminated rocks (rinks dal member) constitute about 13  000 km3 (61%) and the contaminated rocks 8200 km3 (39%) of the succession. within the delimited area, the volume of the rinks dal member is reasonably well constrained. it has a pronounced depocentre west of the disko gneiss ridge where the most complete succession with the oldest part is present and thicknesses exceed 1200 m (fig. 8; central disko section). in the north, the lavas gradually onlapped the shield of the earlier vaigat formation that rose to the north so that only the upper rinks dal member reached far into nuussuaq. the upper rinks dal member originally extended east of 52°w, but little is left of it except for small outliers on eastern nuussuaq. the rinks dal member constitutes between 60% and 66% of the volume of the maligât formation. the volumes of the constituent units of the member have been estimated separately (table 12); with about 7000 km3 the upper rinks dal member constitutes more than half the total volume of the member. the nordfjord member is of large lateral extent (fig. 12) but of limited thickness and volume (1200 km3) and constitutes only around 6% of the maligât formation (table 12). it has a clear depocentre on north-western disko where thicknesses reach 350 m and eruption sites, intermediate lavas and acid tuffs are present (fig. 14). however, over most of the area it is only represented by a few lava flows with combined thicknesses of 30−100 m. the niaqussat member is also of large lateral extent, but its thickness is uncertain because of erosion and removal of much of its upper part. over much of central and western nuussuaq (east of the itilli fault), the member has been completely removed. erosional remnants are present on eastern disko with thicknesses up to 180 m in frederik lange dal where the lower, middle and upper volcanic episode code area (km2) average thickness (m) volume (km3) % of total niaqussat member 530 18000 600–240, av. 400 7000 33 nordfjord member 520 18000 200–30 1200 6 rinks dal member 505–518 18000 1400–200 13000 61 maligât formation, total volume 18000 2200–200 21200 100 units of the rinks dal member upper rinks dal member 514–518 18000 600–200, av. 400 7000 33 akuarut unit 513 13300 350–0, av. 170 2300 11 skarvefjeld unit 511 5000 200–0, av 70 350 2 lower rinks dal member (excl. 511) 505–512 13300 600–0, av. 250 3500 16 volumes are measured within an area bounded by the north coast of nuussuaq, the itilli fault, longitudes 55°w and 52°w, and latitude 69°15′n through qeqertarsuaq town. thicknesses were contoured, sub-areas measured, and volumes calculated in 100-m thick layers. av.: average. table 12. volumes of different parts of the maligât formation 216216 niaqussat member are all present. the best constraint comes from the fact that the upper niaqussat member is present with nearly equal thicknesses (150−190 m) in the most complete profile on western disko (sapernuvik, fig. 14) and in the easternmost volcanic outliers on nuussuaq (point 2080 m, fig. 12). in both profiles, the two uppermost flows even have the same slightly changed geochemical character, suggesting that the succession on eastern nuussuaq is close to the top of the member. in the sapernuvik area, the entire niaqussat member is 500 m thick. the lower niaqussat member seems to be thinner in the east where the whole member may have been around 300 m thick. a good estimate of the average thickness of the niaqussat member is thus 400 m. this means that the niaqussat member originally had a volume of about 7000 km3 and constituted as much as 33% of the maligât formation (table 12). if a thickness of 300 m is used instead, the volume decreases to 5500 km3, i.e. 28% of the maligât formation. in any case, the niaqussat member originally constituted a considerably larger proportion of the maligât formation than it does at present. thickness of the removed succession the average thickness of the niaqussat member as estimated above (400 m) means that a succession of about this thickness has been removed over large areas of disko and nuussuaq where the member is missing. the original extent and thickness of the sapernuvik member is unconstrained but may not have been large. on the other hand, the frequent occurrence throughout disko and nuussuaq of dykes with compositions and ages identical to those of the 60–58 ma svartenhuk formation and the 56–54 ma naqerloq formation (larsen et al. 2016) that form the thick basaltic lava successions on ubekendt ejland and svartenhuk halvø indicate that these lava formations have also been present on disko and nuussuaq. this means that the removed succession was probably significantly thicker than 400 m. uplift studies by japsen et al. (2005, 2009) have indicated that about 1 km of rock succession has been removed from disko and nuussuaq, in good agreement with the above. 217 crustal contamination of the volcanic rocks summary of the main features of the crustal contamination processes • the contaminants are carbon-bearing mudstones and sandstones of the nuussuaq group; these occur abundantly as xenoliths in the contaminated magmas and have the required chemical and isotopic compositions. • degrees of contamination vary from 2−5% in the basalts to 10−50% in the basaltic andesites, andesites, dacites and rhyolites. no rocks more evolved than basalt were produced by ordinary fractional crystallisation. • magma modification was mainly caused by mixing with partial melts from the sediment sidewall and xenoliths, including transfer of sedimentary sulfur and organic carbon to the melt. selective exchange of elements, e.g. ca and fe, also took place. • progressive reduction processes during heating and pressure decrease led to formation of troilite, native iron and graphite. • oxygen fugacities varied by as much as eight orders of magnitude from uncontaminated magmas to xenoliths carrying oxygen-deficient, ti3+-bearing (magnéli) phases. crustal contamination of the nordfjord and niaqussat member magmas is interpreted to have taken place at two different levels: the basalts became slightly contaminated in deep-seated magma chambers in the lower crust, whereas basaltic andesites, andesites, dacites and rhyolites formed in high-level magma chambers by strong contamination with carbon-bearing mudstones and sandstones (larsen & pedersen 2009). a general feature of the strongly contaminated rocks is signs of progressive reduction, delayed or suppressed magnetite crystallisation, unusual mineral zonation patterns, and formation of native iron, sulfides and graphite (e.g. pedersen 1981). below, we discuss the character of the contaminants, the degrees of contamination and the contamination processes for the magmas of both the vaigat and the maligât formations. possible contaminants on their way through the crust, the magmas could have reacted with both the basement gneisses and the up to 6–8 km thick sediments in the nuussuaq basin. sediment xenoliths in more or less modified and re-equilibrated states have been found in the contaminated volcanic units, providing direct evidence of the contamination processes. basement xenoliths are rare, but the possibility of basement contamination cannot be excluded. in the following, compositional data for the possible contaminants are presented in figs 193–195, and the contamination processes and geochemical changes during contamination are summarised. basement the exposed basement in the disko bugt region mainly consists of archaean orthogneisses of tonalitic to trondhjemitic composition that were reworked during the proterozoic; other major archaean units are the atâ tonalite and the rodebay granodiorite (garde & steenfelt 1999). another major unit of the basement is the thick succession of proterozoic metagreywackes in the karrat group that occurs over large areas north of nuussuaq; in the disko–nuussuaq region, it may be present at depth within the substrate of the nuussuaq basin. compositional and isotopic data on the various rock units have been published by kalsbeek et al. (1988, 1998), kalsbeek & skjernaa (1999) and kalsbeek & taylor (1999) but rarely on the same samples. in table 13, we provide systematic compositional and isotopic data for samples of the different basement units. the chemical compositions of the basement samples vary within quite narrow limits. the rodebay granodiorite has distinctly high k2o (3.5 wt%), rb, and ba. the metagreywackes (average of 23 analyses) have higher mgo (3.5 wt%) and lower na2o than the igneous basement rocks, but the trace element compositions are not very different. the multi-element patterns (fig. 195) show a characteristic feature of the metagreywackes: they have higher contents of zr, y and the heavy ree than the igneous rocks, presumably because heavy minerals such as zircon are concentrated in the sediments relative to their igneous sources. sediments the exposed sediments in the nuussuaq basin are fluvial, deltaic and marine, with fluvial facies dominating in the south-east and marine facies dominating in the north30 40 50 60 70 80 90 100 sio2 0 0 0 10 20 30 40 0 2 4 6 8 10 12 14 16 18 20 feo* 2 4 6 8 10 12 14 16 18 0.5 1.0 1.5 2.0 2.5 0 0 1 2 3 4 mgo (wt%) k2o 0.01 0.10 1.00 10.00 mgo (wt%) p2o5 1 2 3 4 5 precambrian gneiss sandstone mudstone sandstone xenolith mudstone xenolith equilibrated sandstone xenolith equilibrated mudstone xenolith 0 2 4 6 8 10 0 2 4 6 8 10 0 2 4 6 8 10 0 2 4 6 8 10 0 2 4 6 8 10 0 2 4 6 8 10 0 2 4 6 8 10 0 2 4 6 8 10 tio2 al2o3 cao na2o sedi-majplot_5_7_16.eps fig. 193. major element variation diagrams for possible contaminants of the volcanic rocks. data in wt% oxides recalculated volatile-free. feo* is total iron as feo. vertical double-headed arrows at the right side of diagrams show the ranges of the respective elements in uncontaminated magmas with 9–18 wt% mgo. 219 west (e.g. g.k. pedersen & pulvertaft 1992). the dominant lithology is arkosic sandstone (schiener 1975). the friable sandstones may be cemented with carbonate, silica or kaolinite. in the fluvial and deltaic successions, the sandstones are interbedded with mudstones and locally with thin coal beds. thicker, non-marine mudstones (the atanikerluk formation, fig. 5) in the south-east have low sulfur contents (up to 0.1 wt% s) and 6−10 wt% toc (g.k. pedersen et al. 1998). marine mudstones (of the itilli and kangilia formations) in the north-west have high sulfur contents (up to 5 wt% s) and mostly 2.5−6 wt% toc (christiansen et al. 1996, 1997; bojesen-koefoed et al. 1997). the gro#3 well on western nuussuaq penetrated c. 3000 m of deep-water marine sediments (dam et al. 2009, fig. 67). the sediments that contaminated the volcanic rocks are situated within the unexposed part of 0 10 20 30 precambrian gneiss sandstone mudstone sandstone xenolith mudstone xenolith equilibrated sandstone xenolith equilibrated mudstone xenolith nb 0 50 100 150 200 ce 0 100 200 300 400 500 600 mgo (wt%) sr 0 100 200 300 400 cu 0 100 200 300 400 ni 0 500 1000 1500 2000 2500 3000 mgo (wt%) cr 0 2 4 6 8 10 0 2 4 6 8 10 0 2 4 6 8 10 0 2 4 6 8 10 0 2 4 6 8 10 0 2 4 6 8 10 sedi-sporplot_6_7_16 fig. 194. selected trace element variation diagrams for possible contaminants of the volcanic rocks. data in ppm. vertical double-headed arrows at the right side of diagrams show the ranges of the respective elements in uncontaminated magmas with 9–18 wt% mgo. 220220 the thick succession but are expected to have lithologies similar to those of the exposed sediments. chemical analyses of mudstones and sandstones from the nuussuaq basin are shown in table 13. these rocks have high contents of volatiles, 10−25 wt% except for a few sandstones, and the recalculation of analyses to 100% volatile-free makes significant differences. the recalculated mudstone analyses form a fairly distinct population in fig. 193. the largest scatter is seen for total iron contents, while a few samples have deviating contents of one or more elements such as na2o, k2o and mgo. the recalculated sandstone analyses show considerable scatter, ranging from a quartzitic sandstone with 90 wt% sio2 to more arkosic varieties with lower 1 10 100 1000 disko gneiss ridge undifferentiated gneiss rodebay granodiorite ataa tonalite karrat group metagreywacke 0 1 10 100 1000 1 10 100 1000 atane fm arkosic sandstone atane fm quartzitic sandstone unequilibrated sandstone xenolith equilibrated sandstone xenolith 0 1 10 100 1000 1 10 100 1000 atane fm mudstone unequilibrated mudstone xenolith equilibrated graphite-rich mudstone xenolith 0 1 10 100 1000 sedi-spid.eps_6_7_16 la ce pr nd pm sm eu gd tb dy ho er tm yb lu rb ba th u nb ta k la ce pb pr sr nd p sm zr hf eu ti gd tb dy y yb la ce pr nd pm sm eu gd tb dy ho er tm yb lu rb ba th u nb ta k la ce pb pr sr nd p sm zr hf eu ti gd tb dy y yb la ce pr nd pm sm eu gd tb dy ho er tm yb lu rb ba th u nb ta k la ce pb pr sr nd p sm zr hf eu ti gd tb dy y yb fig. 195. ree and multi-element diagrams for possible contaminants of the volcanic rocks. left diagram, chondrite normalised; right diagram, primitive mantle normalised; normalisation factors from mcdonough & sun (1995). 221 1001 1001 1001 1001 1001 1001 1001 9202 9102 6274 6274 4204 157205 348667 360907 360994 177375 177378 177379 176769 176770 113449 113450 136992 -5332.66 -5032.26 -5044.15 -5025.39 -5145 -5145 -5145 not def. not def. 5416.36 5416.36 5317.60 6915.94 6949.38 6922.55 7010.89 7115 7115 7115 not def. not def. 6956.19 6956.19 7011.78 176.17 not def. not def. 4.88 4.88 33.25 71.73 69.34 71.43 71.84 59.96 71.58 72.72 44.85 50.09 53.74 51.03 54.30 0.37 0.53 0.2 0.28 0.79 0.60 0.54 1.03 0.97 0.84 0.87 1.21 13.62 15.24 14.94 14.5 15.94 10.96 11.11 23.68 22.9 25.44 25.91 27.80 1.09 3.53 2.03 2.6 0.74 0.54 0.52 0.62 2.28 6.64 4.77 3.55 5.21 4.87 5.63 6.8 7.73 0.04 0.05 0.02 0.03 0.06 0.10 0.08 0.05 0.04 0.06 0.04 0.04 0.99 1.04 0.49 0.67 5.14 3.53 2.96 0.87 0.96 1.45 0.91 1.58 3.22 3.4 2.26 2.66 1.33 5.07 4.79 0.53 0.68 1.13 0.81 1.30 3.9 4.31 4.16 4.38 3.46 0.46 0.34 0.62 0.78 2.3 2.65 1.47 1.25 1.94 3.47 2.03 3.633 1.107 1.734 1.65 2.03 1.7 1.36 1.880 0.08 0.13 0.07 0.08 0.100 0.095 0.086 0.159 0.13 0.14 0.24 0.170 0.44 0.5 0.2 0.3 1.64 0.99 1.39 20.69 12.26 6.7 9.03 1.48 99.01 100.01 99.27 99.37 99.43 99.80 99.82 99.34 99.67 99.75 99.65 99.99 37.5 71.8 62.6 69.1 156 124 94 71.4 68.4 93 110 67 6.3 10.7 4.3 5.3 91.2 11.5 14 42.2 37.7 84 51 67 49.9 57.9 2 4.8 33.3 48.6 40 16.5 14.8 22 34 25 7.5 6.5 7.2 7.8 85.7 53.6 54 50.9 50.1 67 102 72 8.7 4.7 3.6 4.6 23.2 15 15 19.2 14.0 22 25 24 30.2 11.8 19.1 28.2 140 109 108 106 126 244 170 142 7.3 48.1 12.8 20.5 195 217 177 134 131 156 177 163 13 20.8 17.8 16.7 24 17.2 17 31.2 30.5 33 38 46.0 50.4 104 97.5 121 88.8 84 64.7 77.3 82.8 57.4 102 238 464 384 139 86.8 93.2 123 201 182 213 193 171 12.9 14.1 12.8 7.76 23.2 21 18 24.6 19.7 25.0 27.1 36 114 104 137 152 133 169 135 173 170 232 134 177 3.01 5.99 5.77 6.54 10.6 10.5 7.9 16.4 13.5 13.8 15.1 17 1.38 1.64 0.84 1.70 6.20 6.00 4.70 3.00 2.87 1.67 2.12 432 543 1208 288 495 392 366 464 490 471 353 478 30.4 18.9 39.3 28.1 29.7 21.1 19.2 53.1 42.7 44.2 47.5 60.6 61.6 46.0 72.6 49.4 56 42 39 111 88.3 105 125 126 7.56 6.22 8.14 6.08 12.5 10.4 11.1 12.8 27.1 23.7 26.2 19.9 23 19 16 45.7 37.7 41.5 47.9 53 4.79 5.20 4.29 3.77 4.03 3.41 3.03 8.13 6.31 7.77 8.31 0.907 1.16 0.652 0.529 0.9 0.78 0.72 1.74 1.35 1.51 1.69 4.39 3.97 3.26 3.11 8.01 6.08 7.57 8.61 0.579 0.543 0.445 0.382 0.5 0.5 0.5 0.938 0.689 0.964 0.995 2.90 2.88 2.38 1.77 4.80 3.66 5.09 5.43 0.512 0.484 0.423 0.266 0.863 0.685 0.923 0.961 1.25 1.30 1.18 0.691 2.42 1.95 2.62 2.70 0.152 0.177 0.170 0.084 0.321 0.280 0.353 0.408 0.918 1.08 1.00 0.501 1.71 2.12 1.67 2.04 1.83 2.25 2.55 0.129 0.154 0.141 0.083 0.24 0.31 0.23 0.297 0.272 0.325 0.358 3.18 2.75 3.50 3.93 2.8 3.8 3.1 4.24 4.16 6.29 3.63 0.181 0.699 0.651 0.422 0.8 0.6 0.5 0.981 0.875 0.834 1.11 5.53 11.3 16.3 14.7 37.8 17.9 14 18.7 21.9 25.0 16.7 6.96 6.77 12.0 8.87 11.9 7.9 7 10.4 8.67 12.9 14.3 15.2 0.411 2.06 0.669 0.651 2.5 1.2 1.1 3.77 2.60 3.22 5.18 0.721167 0.713490 0.732382 0.781592 0.81954 0.76145 0.76212 0.725067 0.720346 0.730705 235.85 126.88 395.03 1093.48 1632.15 807.55 817.05 291.20 224.19 371.22 0.511068 0.511221 0.510750 0.510909 0.511270 0.511475 0.511469 0.511121 0.511123 -29.12 -26.14 -35.32 -32.22 -25.19 -21.19 -21.30 -28.10 -28.05 16.948 18.581 14.964 14.878 17.364 18.099 18.878 17.692 17.526 15.189 15.389 14.705 14.719 15.493 15.575 15.712 15.192 15.135 42.659 37.540 39.379 38.261 36.948 37.681 38.334 37.966 37.873 lith. code ggu no. deg. w deg. n altitude, m sio2 tio2 al2o3 fe2o3 feo mno mgo cao na2o k2o p2o5 volatiles sum zn cu co ni sc v cr ga rb sr y zr nb cs ba la ce pr nd sm eu gd tb dy ho er tm yb lu hf ta pb th u isotope ratios calculated at 60 ma 87sr/86sr60 εsr 143nd/144nd60 εnd 206pb/204pb60 207pb/204pb60 208pb/204pb60 for explanation of lithological codes, see table 1. for petrographical notes on the samples, see table 13c. geographical coordinates in wgs 84. first two digits are degrees, then follow minutes in decimal form. not def.: not definable. major elements in wt% (xrf analyses). trace elements in ppm (zn–ga: xrf and some icp-ms analyses; rb–u: icp-ms analyses). trace elements in 177375-379: xrf and ina analyses; in samples 136992, 113306 and 176506 only xrf analyses. high contents of analytical volatiles are due to high contents of toc and s (see table 13c). mudstoneprecambrian basementlithology table 13a. chemical analyses of precambrian basement, nuussuaq basin sediments, and xenoliths 2201 4204 2201 2201 2201 4204 6274 6284 4204 5207 5204 113202 113223 113493 176771 176772 113306 113443 113527 400296 156518.2 176506 5332.67 5332.79 5314.20 5315 5300.0 5317.69 5416.36 5411.94 5330.08 5446.31 5442.88 7013.42 7012.92 7010.76 7011 7004.5 7011.81 6956.19 6921.64 7029.95 7007.633 7014.60 360.19 535.62 3.67 25.64 4.88 1144.77 160 420.21 73.29 88.97 87.77 53.61 55.15 46.08 81.84 32.01 45.07 52.47 66.19 0.75 0.06 0.05 0.71 1.31 0.27 0.33 0.56 0.54 0.57 0.24 5.15 2.24 5.82 10.53 11.31 25.29 9.1 24.83 26.52 16.97 7.46 0.08 1.50 0.92 0.34 0.51 3.40 1.76 0.92 1.29 0.48 0.2 1.65 4.06 2.74 1.35 8.15 0.00 1.37 1.87 0.02 0.03 0.02 0.1 0.09 0.06 0.04 0.09 0.04 0.03 2.62 2.08 0.28 0.77 3.63 2.9 0.83 4.46 3.74 0.92 6.51 5.32 0.85 0.71 13.73 6.58 9.02 3.12 7.82 7.67 1.64 11.95 0.81 0.41 0.69 1.63 0.56 1.41 1.03 0.92 0.77 3.26 0.99 2.12 0.260 2.08 2.28 2.34 0.59 0.45 0.080 0.408 6.150 1.670 0.06 0.040 0.07 0.07 0.01 0.060 0.030 0.120 0.020 8.09 2.27 1.89 14.05 14.67 10.47 1.38 19.63 11.69 14.47 2.09 99.60 99.19 99.51 99.98 100.11 98.93 99.99 98.61 99.87 99.70 99.94 21.0 8.32 7.41 33.9 53.7 63 21.6 126 88.1 40.3 8 2.94 10.4 1.45 7.42 10.7 65 7.18 227 29.7 22.0 12 4.30 18.9 1.77 8.85 8.98 15 3.12 56.3 14.9 6.32 4 11.0 30.0 4.73 17.4 23.7 94 12.1 287 47.4 13.9 17 8.61 3.93 1.11 8.18 16.2 8.3 5.67 11.8 8.93 9.57 4.6 62.6 227 6.40 32.5 63.1 176 26.7 318 132 72.6 29 58.5 34.5 11.7 54.4 78.6 910 33.9 870 569 77.8 35 5.64 3.39 5.65 10.9 14.5 9.94 26.9 29.0 23.8 9 48.6 11.2 47.9 58.3 57.7 18 5.26 2.31 14.32 83.4 43 382 92.3 83.7 1075 115 342 245 285 298 379 244 25.1 3.40 2.54 14.1 22.8 6 8.43 8.18 8.04 32.4 4.9 657 28.4 40.4 294 711 62 184 42.0 123 17.9 311 9.94 0.51 1.16 9.86 17.6 3 5.25 1.72 6.67 8.55 2.6 0.388 0.166 0.308 0.886 0.809 0.122 0.115 0.467 0.448 569 183 484 705 598 284 363 59.32 219 2653 554 41.5 3.80 3.64 20.9 38.0 10 11.7 7.551 17.5 48.7 9.8 78.3 6.91 6.61 38.4 72.9 19 21.3 14.5 33.2 93.2 20 8.83 0.791 0.796 4.44 8.42 2.58 1.71 3.79 10.7 30.5 2.91 2.81 15.9 29.7 7 9.13 6.95 13.7 38.8 11 4.78 0.531 0.482 2.82 4.98 1.70 1.62 2.50 7.24 0.577 0.103 0.174 0.572 0.723 0.595 1.34 1.29 2.74 5.27 0.590 0.487 3.08 5.38 1.95 1.75 2.23 6.72 0.627 0.082 0.078 0.392 0.619 0.269 0.264 0.282 1.02 3.86 0.495 0.438 2.31 3.46 1.53 1.50 1.58 5.49 0.808 0.099 0.087 0.450 0.731 0.307 0.301 0.319 1.03 2.50 0.275 0.245 1.34 2.22 0.887 0.815 0.896 2.71 0.399 0.039 0.033 0.207 0.356 0.135 0.108 0.125 0.390 2.69 0.223 0.206 1.30 2.35 0.863 0.688 0.834 2.21 0.419 0.031 0.033 0.209 0.375 0.130 0.095 0.116 0.321 14.9 0.639 1.16 7.42 18.2 4.40 1.19 3.64 0.818 0.677 0.068 0.083 0.579 1.13 0.324 0.222 0.558 0.568 10.6 0.398 8.223 12.7 12.3 4.60 6.066 3.80 46.0 6 13.5 0.186 0.831 6.25 14.9 1.95 3.12 0.833 5.70 7.87 4 2.17 0.065 0.256 1.53 3.20 0.837 0.193 1.04 1.43 0.718483 0.718084 0.717170 0.717855 0.714886 0.727655 197.75 192.09 179.12 188.84 146.70 327.93 0.510700 0.510962 0.511028 0.511890 -36.30 -31.19 -29.91 -13.09 17.832 16.316 17.226 16.973 15.205 14.954 15.118 15.037 46.498 37.227 37.309 37.297 lith. code ggu no. deg. w deg. n altitude, m sio2 tio2 al2o3 fe2o3 feo mno mgo cao na2o k2o p2o5 volatiles sum zn cu co ni sc v cr ga rb sr y zr nb cs ba la ce pr nd sm eu gd tb dy ho er tm yb lu hf ta pb th u isotope ratios calculated at 60 ma 87sr/86sr60 εsr 143nd/144nd60 εnd 206pb/204pb60 207pb/204pb60 208pb/204pb60 magma-equilibrated sediment xenoliths buchitesandstonelithology table 13b. chemical analyses of precambrian basement, nuussuaq basin sediments, and xenoliths 222 223 sio2. the sandstones with more than 5 wt% cao, three of which also have high mgo (3−4 wt%), are carbonatecemented. some of the sandstone samples have very high contents of zr (up to 1650 ppm recalculated), indicating high contents of detrital zircon grains. elements such as y and hree, which are concentrated in zircon, are therefore also quite high. the arkosic sandstones have lree contents similar to those of the basement gneisses but significantly increased hree and y contents due to relative accumulation of heavy minerals such as zircon (fig. 195). this ’heavy mineral effect’ is more pronounced in the mature sandstones than in the mudstones and the immature karrat group metagreywackes. except for this effect, the multi-element patterns of the sediments are largely similar to those of the meta-igneous basement rocks from which the sediments were ultimately derived. thus, all samples have high contents of rb-ba-th-u and k-la-ce, and significant nb-ta troughs, pb peaks, and p and ti troughs in the multi-element patterns. distinction between contaminants as discussed by larsen & pedersen (2009), the sr, nd and pb isotope data enable us to distinguish between the different potential contaminants. the proterozoic metagreywackes of the karrat group have higher 207pb/204pb ratios (15.5–15.7) than both the uncontaminated and contaminated magmas (207pb/204pb = 15.2–15.5 and 14.9–15.4, respectively), thereby excluding any significant contamination from these rocks. the four local basement orthogneisses analysed in this work have quite variable isotope ratios, and so have other orthogneisses from the region (kalsbeek & taylor 1999) which have 87sr/86sr up to >1. in contrast, the sediments of the nuussuaq basin have restricted sr isotopic compositions, which are within the middle range of the orthogneisses, as expected for sediments derived from these gneisses (fig. 196). in all, the sediments serve very well as possible contaminants, as detailed below. contamination by basement orthogneisses with relatively low 87sr/86sr cannot be excluded but is not necessary to explain the data. 157205 gneiss c. 6 m below the contact to the volcanic rocks, disko gneiss ridge below lyngmarksfjeld, southern disko. 348667 atâ tonalite, illuluarsuit nunataat island east of northern arveprinsen ejland. 360907 rodebay granodiorite, paakitsup nunaa c. 25 km north-north-east of ilulissat. 360994 undifferentiated gneiss, easternmost nuussuaq. 177375 karrat group metagreywacke, alfred wegener halvø, uummannaq fjord. 177378 karrat group metagreywacke, alfred wegener halvø, uummannaq fjord. 177379 karrat group metagreywacke, alfred wegener halvø, uummannaq fjord. 176769 composite of 12 unmetamorphosed cretaceous to paleocene mudstone samples from disko. 176770 composite of 11 unmetamorphosed cretaceous to paleocene mudstone samples from nuussuaq. pedersen (1979a), new icp-ms data. ‘volatiles’ include 5.43 wt% toc and 0.85 wt% s. 113449 mudstone xenolith, unequilibrated, in nordfjord member fe-andesite dyke, innermost nordfjord, western disko. ‘volatiles’ include 1.86 wt% toc and 2.46 wt% s. 113450 mudstone xenolith, graphite-rich, unequilibrated, in nordfjord member fe-andesite dyke, innermost nordfjord, western disko. pedersen & larsen (2006), new icp-ms data. ‘volatiles’ include 4.07 wt% toc and 3.05 wt% s. 136992 mudstone xenolith, slightly equilibrated, in fe-andesite lava flow of the asuk member at asuk, northern disko. pedersen (1979a). ‘volatiles’ include 0.33 wt% toc and 0.78 wt% s. 113202 cretaceous sandstone, maniillat kussinersuat gully, northern disko. 113223 sandstone xenolith in basaltic andesite lava flow of the asuk member, maniillat kussinersuat gully, northern disko. 113493 cretaceous sandstone, quartz-rich, coast south-east of asuk, northern disko. 176771 cretaceous sandstone with muddy streaks, asuk. sample collected by k.j.v. steenstrup 1872. 176772 cretaceous sandstone with plant imprints, ritenbenks kulbrud south of qullissat. sample collected by k.l. giesecke july 17–18 1811 (217 no. 235). 113306 mudstone xenolith, graphite-rich, equilibrated, in fe-andesite lava flow from asuk member at asuk, northern disko. pedersen (1979a). ‘volatiles’ include 6.98 wt% toc and 0.06 wt% s. 113443 sandstone xenolith, glass-rich, equilibrated, in nordfjord member fe-andesite dyke, innermost nordfjord, western disko. 113527 mudstone xenolith, graphite-rich, equilibrated, from dyke on uiffaq, south-west disko. sample collected by k.j.v. steenstrup 1881. 400296 mudstone xenolith, graphite-rich, equilibrated, in graphite andesite tuff at ilugissoq, central nuussuaq. pedersen & larsen (2006), new icp-ms data. ‘volatiles’ include 8.83 wt% toc and 0.03 wt% s. 156518.2 mudstone xenolith, k2o-rich and graphite-rich, equilibrated, in garnet rhyolite lava block in conglomerate. sedimentkløften, south side of hammer dal, north-west disko. toc and s in table 7. 176506 sandstone buchite xenolith in fe-andesite lava flow of the nordfjord member at toornivit, north-west coast of disko. table 13c. notes on analysed samples of precambrian basement, sediments of the nuussuaq basin, and xenoliths 224224 contamination by an enriched lithospheric component in some rocks was identified by larsen et al. (2003) and larsen & pedersen (2009). this took place at deep levels before the magmas entered high-level magma chambers and reacted with sediments and is not considered further here. degrees of contamination as shown in fig. 196 the total sr-nd isotope data set conforms well to a simple model of bulk mixing between a picrite and an average sediment. the various contaminated units of both the vaigat and maligât formations are shown in details in fig. 197a. the major contaminated members of the vaigat formation require more than about 10% contamination except for the nuusap qaqqarsua member which only requires 5−10% and whose chemical composition is also less modified (pedersen et al. 2017, fig. 156). the asuk member rocks are contaminated with up to 40% sediment, which is mudstone-dominated according to pedersen & pedersen (1987) and goodrich & patchett (1991). the kûgánguaq member is contaminated with 12–24% sediment, which is sandstone-dominated according to pedersen & pedersen (1987). in contrast the contaminated members of the maligât formation (nordfjord and niaqussat members) represent a large range of contamination from only 12% to >50% and c. 35%, respectively, in accordance with their large compositional ranges. the most highly contaminated rocks of all are the rhyolites of the nordfjord member, which contain around 50% sediment, which is mudstone-dominated according to pedersen & pedersen (1987). very slight contamination may be difficult to detect. some samples that are considered contaminated because of small compositional anomalies appear to contain only around 1% contaminant (fig. 197b). on the other hand, three samples from the low-ti unit in the anaanaa member have sr-nd isotope ratios that suggest 2−3% contamination even though their chemical compositions do not match this. this unit is possibly not contaminated, as discussed by larsen & pedersen (2009). in general, we consider a sample to be crustally contaminated if it has 0.5106 0.5111 0.5116 0.5121 0.5126 0.5131 0.70 0.712 0.722 0.732 0.742 0.752 0.762 87sr/86sr(60) 14 3 n d/ 14 4 n d( 60 ) uncontaminated contaminated alkaline/enriched contaminated alkaline/enriched basement gneiss basement metasediment sandstone equilibrated sandstone xenolith shale equilibrated shale xenolith b a s e m e n t o r t h o g n e i s s e s 40% 20% 10% 5% bulletin_jp16bfig. 196. all sr and nd isotope data for the vaigat and maligât formations and some possible contaminants (slightly modified from larsen & pedersen 2009). four basement orthogneisses are from nuussuaq, disko and the mainland east of disko (one sample is off scale with 87sr/86sr(60) = 0.7816). other local basement orthogneisses with 87sr/86sr(60) up to >1 are from kalsbeek & taylor (1999). the basement metasediments are from the proterozoic karrat group (kalsbeek et al. 1998). the shales and sandstones are cretaceous sediments from the nuussuaq basin. the curve is a mixing curve calculated by bulk mixing of a picrite with 120 ppm sr, 7 ppm nd, 87sr/86sr(60) = 0.7031, and 143nd/144nd(60) = 0.5130 with an average sediment with 230 ppm sr, 39 ppm nd, 87sr/86sr(60) = 0.7250, and 143nd/144nd(60) = 0.5111. tick marks indicate the amount of sediment (wt%) in the mixture. 225 several or all the following characteristics relative to ‘normal’ rocks: increased sio2, rb, ba, th, k, lree, pb, and 87sr/86sr, and decreased 143nd/144nd. nb and zr may not decrease and there need not be any nb or zr trough in the multi-element pattern because some sediments have quite high nb and zr (up to 17 ppm nb and 1100 ppm zr). an important conclusion of pedersen & pedersen (1987) is that no rocks more evolved than basalt were produced by ordinary fractional crystallisation; this is probably the result of very high magma production and throughput rates. more evolved rocks are always contaminated and were produced in local magma chambers with limited or no input of fresh magma. in such chambers the most evolved magmas that were formed were highsi rhyolites which had fractionated plagioclase, sanidine, ti-oxide and zircon, judged from the sr, ba, eu, ti and zr troughs in the multi-element patterns (fig. 148). the minimum mgo content in the parent magma can be estimated from the cr content in the contaminated magma because chromite crystallisation was delayed by low fo2 caused by organic compounds from the sediments (pedersen 1985). contamination processes and geochemial changes during contamination the contamination processes in the magma chambers are very complex and have been dealt with in numerous papers, viz. melson & switzer (1966), pedersen (1978a,b, 1979a,b, 1981, 1985), ulff-møller (1979, 1985, 1990), pedersen & pedersen (1987), pedersen & rønsbo (1987), goodrich & patchett (1991), lightfoot et al. (1997) and larsen & pedersen (2009). sediment xenoliths in the volcanic rocks provide direct evidence for the contamination processes, and analyses of slightly to strongly magma-modified and equilibrated xenoliths are included in table 13. mixing and afc processes bulk mixing of disintegrated sediment fragments into the magma took place to some extent, but the major magma modification was caused by mixing with partial melts of the sediment sidewall in high-level magma chambers. as evidenced by pyrometamorphosed and melted xenoliths (buchites), the high temperatures of the magmas induced partial melting of the sidewall, including the loosened xenoliths, and it was these melts, which were of rhyolitic composition, that mixed with the magmas. analyses of melts from buchites are presented in table 14. on a volfig. 197. a: degrees of crustal contamination in the various contaminated units of the vaigat and maligât formations. the mixing model is the same as shown in fig. 196. tick marks indicate the amount of contaminant (wt%) in the mixture. a model using a parent basalt with slightly higher nd and sr instead of a picrite would not change the mixing curvature significantly but would shift the tick marks upwards to require slightly more contaminant. b: enlargement of the less contaminated part of a. slightly modified from larsen & pedersen (2009, fig. 19). 0.5114 0.5116 0.5118 0.5120 0.5122 0.5124 0.5126 0.5128 0.5130 0.5132 uncontaminated niaqussat mb nordfjord mb contam. enriched/alkaline enriched/alkaline manîtdlat mb contam. ordlingassoq tunoqqu mb kûgánguaq mb asuk mb nuuk killeq mb nuusap qaqqarsua mb marraat unit other contaminated 40% 30% 20% 10% 5% 50%a 0.5125 0.5127 0.5129 0.5131 uncontaminated contaminated alkaline/enriched 4% b 1%0.5% 2% 3% 5% 0.702 0.702 0.703 0.704 0.705 0.7060.706 0.710 0.714 0.718 14 3 n d/ 14 4 n d( 60 ) 14 3 n d/ 14 4 n d( 60 ) 87sr/86sr(60) 87sr/86sr(60) bulletin_jp19 226226 atile-free basis, these melts contain 72−77 wt% sio2 and 6−8 wt% alkalis with high k/na ratios. the mudstone buchite melts are peraluminous. the calculated result of simple mixing of such a rhyolitic buchite melt into a picrite magma with 15 wt% mgo is shown in table 15. in the contaminated melt, sio2 and k2o are increased, feo*, mgo and cao are decreased, and tio2, al2o3, na2o and p2o5 are nearly unaffected. such changes can be seen in all the contaminated rocks in west greenland. the calculated mixture with 10% crustal melt is actually very close in composition to the magmas of the nuusap qaqqarsua member of the vaigat formation. if the contaminated magmas also fractionate, sio2 and k2o will increase further and mgo and cao will decrease further, whereas the lowered tio2 and feo* will be counteracted by fractionation-induced increases. assimilation coupled with fractional crystallisation (afc) was modelled with sr isotopes for the contaminated magmas in west greenland by pedersen & pedersen (1987), and bulk mixing was modelled with sr and nd isotopes by goodrich & patchett (1991), and with trace elements by lightfoot et al. (1997). the isotopes are easily modelled but do not yield much information about other processes, such as assimilation of partial melts and exchange of other elements. reduction processes when sediments are heated by magmas, they lose volatile components to the melt, mainly water but also sulfur and carbon compounds. if the sediments are rich in sulfur and carbon, a chain of p−t-dependent reduction processes takes place in the magma that may ultimately lead to the formation of sulfides, native iron (alloyed with carbon) and graphite. the c−co−co2 buffer equilibrium plays a key role in the processes (e.g. pedersen 1981). this equilibrium is strongly pressure-dependent (french & eugster 1965; french 1966; sato 1978); at pressures very much higher than atmospheric, such as prevails in a magma chamber, the magma may not be very reduced, but by pressure release during high-level intrusion or surface eruption large amounts of oxygen are consumed by this reaction, and liquid sulfide and native iron may form. the contaminated rocks of the vaigat and maligât formations and their xenoliths display a range of oxides, metals and sulfides which indicate that the oxygen fugacities at high temperatures varied by as much as eight orders of magnitude (fig. 198) from uncontaminated picrites 113449 113450 176522 138237 176506 average glass mudstone mudstone mudstone sandstone sandstone 6 1σ 8 1σ 3 1σ 2 1σ 7 1σ 26 recalculated 67.94 0.49 66.45 0.52 69.40 0.82 69.20 0.14 73.82 0.86 69.36 73.89 1.00 0.12 1.23 0.06 1.73 0.31 0.45 0.03 0.26 0.06 0.94 1.00 14.79 0.53 15.59 0.23 13.53 1.10 13.35 0.21 11.62 0.25 13.78 14.70 0.82 0.09 0.25 0.05 2.90 0.87 1.75 0.07 0.14 0.11 1.17 1.25 0.02 0.02 0.54 0.09 0.81 0.10 0.94 0.35 1.30 0.00 0.25 0.14 0.77 0.82 1.73 0.10 1.27 0.06 0.95 0.65 1.35 0.07 1.83 0.21 1.43 1.52 3.03 0.34 3.69 0.95 3.23 0.67 2.85 0.07 2.01 0.11 2.96 3.16 3.57 0.20 2.40 0.07 2.40 0.36 2.55 0.49 5.84 0.15 3.35 3.56 0.48 0.05 93.41 92.18 95.09 92.80 95.78 94.24 100.00 ggu no. xenolith type n sio2 tio2 al2o3 feo mno mgo cao na2o k2o p2o5 sum n: number of analyses table 14. microprobe analyses of glasses in partially melted sediment xenoliths (buchites) picrite melt 9:1 mix difference relative 400176 difference, % 47.5 73.9 50.14 2.64 5.6 49.73 1.2 1 1.18 -0.02 -1.7 1.2 12.5 14.7 12.72 0.22 1.8 12.95 11.5 1.3 10.48 -1.02 -8.9 10.68 15 0.8 13.58 -1.42 -9.5 13.11 10.5 1.5 9.6 -0.9 -8.6 9.92 1.5 3.1 1.66 0.16 10.7 1.84 0.1 3.6 0.45 0.35 350 0.37 0.1 0.2 0.11 0.01 10 0.1 99.9 100.1 99.92 99.9 sio2 tio2 al2o3 feo mgo cao na2o k2o p2o5 sum picrite: simplified typical composition, naujánguit member. melt: average of five buchite glasses (microprobe analyses), table 14. difference: difference between between mixture and picrite. 400176: contaminated lava flow, nuusap qaqqarsua member. compare 400176 with the calculated mixture. table 15. contamination of picrite by simple mixing with rhyolitic melted sediment 227 -3 -4 -5 -6 -7 -8 -9 -10 no n h m mq f m w w fe llm fe-f b llm feru ti 6o 11 ti 5o 9 ti 5 o 9 ti 4 o 7 ti 4o 7 ti 3o 5 ti 3o 5 ti 2o 3 ti 2o 3 tio ti 7o 13 ti 6o 11 f b fe-ru -11 -12 -13 -14 -15 -16 -17 -18 -19 -20 -21 1000 1100 1200 t°c gv03_03_068_lml log fo2 1 2 4 5 7 8 6 3 fig. 198. fo2–t diagram at 1 bar pressure for a number of investigated volcanic rocks and their xenoliths from the vaigat and maligât formations on disko and nuussuaq. a series of oxygen buffer curves are shown as reference curves, with the following abbreviations: h: hematite. m: magnetite. no: nickel oxide. n: nickel metal. q: quartz. f: fayalite. w: wüstite. fe: metallic iron. ilm: ilmenite. fb: ferropseudobrookite. ru: rutile. the stabilities of a number of reduced ti-oxides are shown with formulae. field 1: uncontaminated volcanic rocks of the vaigat formation (pedersen 1985). field 2: kûgánguaq member, disko. magnesian basaltic andesite and magnesian andesite lava flows with very low fe3+ but no native iron and only trace amounts of sulfides; anomalously low cu and ni indicate previous sulfide fractionation (pedersen 1985). field 3: killiit dyke, disko. chilled margin of magnesian basaltic andesite glass with native iron and sulfide globules (pedersen 1979b). field 4: asuk member tuffs in agatdalen, nuussuaq. graphite-rich magnesian andesites with sulfide globules with wüstite and iron metal (pedersen 1978b). field 5: andesite and dacite lava flows, nordfjord member, disko, with ilmenite phenocrysts reacted to rutile, iron metal and armalcolite (pedersen 1981). field 6: xenolith in native-iron-bearing lava flow, asuk member, disko. partially melted and chilled mudstone containing native iron, troilite, rutile and armalcolite (pedersen 1979a). field 7: magnesian andesite lava flow, asuk member, disko, with native iron and rutile (unpublished). field 8: xenolith in nativeiron-bearing magnesian andesite lava flow, asuk member, disko. heated and chilled mudstone with oxygen-deficient ti-oxides (pedersen & rønsbo 1987). 228228 and basalts to xenoliths carrying ti3+ -bearing, oxygendeficient magnéli phases (general formula tino2n−1, e.g. pedersen 1978b, 1979a, 1981, 1985; pedersen & rønsbo 1987). the formation of native iron can readily be explained by well-known terrestrial reduction processes involving carbon components from sediments, and there is no need to appeal to exotic sources such as earth’s sublithospheric mantle (bird & weathers 1977) or iron meteorites (nordenskiöld 1871; jones et al. 2005). exchange of elements major elements. very large decreases in feo* (and ni, cu and pge) as seen in the asuk and kûgánguaq member rocks are caused by fractionation of sulfides or native iron. the large decreases in cao and mgo in many of the contaminated rocks are caused by loss of these elements from the magma into the sidewall. evidence for this is found in the magma-equilibrated mudstone xenoliths: the highly aluminous mudstones have scavenged cao and mgo from the magma to form calcic plagioclase and magnesian spinel (melson & switzer 1966; pedersen 1978a). the completely recrystallised and equilibrated xenoliths consist of plagioclase-spinel-corundum-graphite aggregates. comparison with the exposed sediments (table 13; fig. 193) shows that the equilibrated mudstone xenoliths have lost feo, tio2, k2o and p2o5, and have gained cao (in gross amounts), mgo and na2o. even the allegedly unequilibrated mudstone xenoliths seem to have gained na2o. na2o shows that an element may participate in counteracting processes: it is both delivered into the magma with the partial melt and scavenged back again through mineral equilibration. other examples of complex element behaviour are the losses of feo* and tio2 in the equilibrated xenoliths. trace elements. figure 195 shows that the equilibrated xenoliths have suffered losses of almost the whole range of incompatible trace elements but have gained sr, which is scavenged with ca into plagioclase. eu is not lost because it is retained as eu2+ in the new-formed plagioclase in the xenolith; the rhyolitic partial melt from the xenolith would have had a pronounced negative eu anomaly without having fractionated plagioclase. the exchange of incompatible elements between mudstone and picrite magma is illustrated in fig. 199. the compositional differences between picrite and mudstone are around two orders of magnitude for the most incompatible elements rb−u (e.g. 0.04 ppm u in picrite, 4 ppm u in mudstone), declining to around one order of magnitude for pr, and to a factor of <2 for dy−yb. the differences between picrite and a rhyolitic partial melt would be even greater. accordingly, the contaminated magma (a magnesian andesite) has very strongly increased concentrations of the most incompatible elements and almost unchanged concentrations of the least incompatible elements. the nb−ta, sr, p and ti troughs and th−u and pb peaks of the sediment pattern are transferred to the pattern of the contaminated magma. it should be noted that despite the nb−ta, sr and p troughs the contami1 10 100 1000 mudstone equilibrated xenolith contaminated magma uncontaminated picrite 1 10 100 1000 sediment–magma reaction uko+kon+sed+xen_bull_6_7_16 la ce pr nd pm sm eu gd tb dy ho er tm yb lu rb ba th u nb ta k la ce pb pr sr nd p sm zr hf eu ti gd tb dy y yb fig. 199. ree and multi-element diagrams illustrating the behaviour of incompatible elements during crustal contamination. left diagram: chondrite normalised; right diagram: primitive mantle normalised; normalisation factors from mcdonough & sun (1995). the uncontaminated picrite is of the naujánguit member (400139, 16.4 wt% mgo), the contaminated magma is an andesite of the kûgánguaq member (135927, 8.5 wt% mgo, 58.4 wt% sio2), the equilibrated mudstone xenolith is of the nordfjord member (400296) and the unmodified mudstone is a composite of 10 mudstone samples from disko (176769). note selective retention of eu and sr in the equilibrated xenolith. see text for discussion. slightly modified from larsen & pedersen (2009, fig. 20). 229 nated magma has not lost these elements, only gained less of them than of the neighbouring elements in the pattern. the eu trough in the contaminated magma is caused by fixation of eu2+ in the sediment by crystallisation of plagioclase, while the other ree were transferred to the melt. in the case of pb, a trough in the uncontaminated picrite with 0.5 ppm pb is turned into a peak in the contaminated magma with 9.5 ppm pb, illustrating that most of the pb in the contaminated magma comes from the sediment. the variation of transition elements in the contaminated rocks (fig. 147; pedersen et al. 2017, fig. 158) is caused by several mechanisms. ni, cu and platinumgroup elements (pge) partition extremely efficiently into sulfide and particularly metal liquid, and a hallmark of sulfide or metal fractionation from a silicate magma is its extremely low contents of ni and cu, as seen in the asuk and kûgánguaq member rocks. some of the ni and cu-depleted rocks carry native iron but other, such as all of the kûgánguaq member rocks, do not. this indicates that sulfide or metal fractionation has taken place at depth before eruption (pedersen 1985; lightfoot et al. 1997). andesites of the asuk member do not show ni and cu depletion, presumably because the andesite magma quickly became too viscous to allow metal and sulfide to settle out. relative enrichment in cr compared to uncontaminated rocks with similar mgo took place because carbon-induced reduction delayed or prevented precipitation of chromite in the melt (pedersen 1985). except for the effects of native-iron and sulfide accumulation or fractionation, the native-iron-bearing contaminated rocks do not show marked compositional differences from the native-iron-free counterparts. in detail, however, the segregated metal-rich bodies show very complex and unusual mineralogies (e.g. goodrich 1984; ulff-møller 1985, 1989). in rare cases, highly reduced fluids have been preserved as inclusions in olivine in basaltic glass chills and have precipitated hydrocarbon minerals there (solovova et al. 2002). 230230 concluding remarks the deposition of the maligât formation succeeded that of the vaigat formation with little or no time lapse between them, and perhaps they even overlapped in time. the transition was presumably caused by changed tectonic conditions, which resulted in establishment of long-lived, deep-seated magma chambers where the picritic primary magmas, which were still produced (larsen & pedersen 2009), were stalled and fractionated to basalt. in this respect, the maligât formation is a more normal volcanic succession than the unique, picritic vaigat formation. the rinks dal member magmas represent conditions of quasi-equilibrium in the deep magma chambers, although periods with variable magma residence times may be inferred from the oscillating degrees of fractionation. the nordfjord member may represent a period with reduced magma production and therefore greater likelihood for magmas to stall at high crustal levels. the niaqussat member reflects a new episode of increased magma production with formation of picrites at its onset but gradually waning to end with quite fractionated basalt magmas. this would have been a suitable end to the magmatic history of the maligât formation; however, the existence of the few lava flows of the sapernuvik member with a suggestion of remelted gabbro at depth shows that a full record of the magmatic history of the formation is probably not preserved. the maligât formation is confined to the southern part of the nuussuaq basin. the contemporaneous evolution in the northern part of the basin (svartenhuk halvø) probably did not comprise deposition of volcanic rocks (any age difference between the uppermost vaigat formation in the south and in the north is unresolved), but a non-marine sediment horizon of sandstones and mudstones with coal seams was deposited ( j.g. larsen & grocott 1991; j.g. larsen & pulvertaft 2000). the volcanic activity in the southern part of the nuussuaq basin did not terminate with the deposition of the maligât formation. following this, two major volcanic episodes led to the deposition of the late paleocene basalts of the svartenhuk formation and the early eocene basalts of the naqerloq formation, which were both originally present over the whole nuussuaq basin from south to north (larsen et al. 2016). these formations contain mass flows and acid tuffs and will be described elsewhere. acknowledgements the results described in this bulletin are based on field work carried out under the auspices of ggu and subsequently geus, and with support from the bureau of minerals and petroleum, government of greenland. arktisk station (university of copenhagen) in qeqertarsuaq also provided generous field support. the work was begun in 1968 (akp) at the suggestion of professor arne noe-nygaard, and over the years we have been indebted to many persons, in particular to the expedition leaders gilroy henderson, feiko kalsbeek and flemming getreuer christiansen, as well as to department leaders niels henriksen, christian knudsen and karen hanghøj for general geus support. the photogrammetric work was mainly carried out at the technical university of denmark in close cooperation with keld s. dueholm and with generous support from ole mærsk-møller and ole jacobi. we are also grateful for support from hans jepsen and erik vest sørensen at the ggu/geus photogrammetric laboratory, to jakob lautrop for reproduction of countless photographs in the pre-digital age and to annette hindø and willy weng for technical and geodetic help in the design and production of the geological maps and sections. jørgen kystol and ib sørensen, and in later years olga nielsen, at ggu/geus’s rock geochemical laboratory, and john bailey at the xrf laboratory of the geological institute, university of copenhagen, maintained the constantly high quality of the chemical analyses of the rocks. jørgen bojesen-koefoed is thanked for the toc and sulfur analyses. jette halskov prepared the drawings and annotated the photographs, and without her patient and excellent help over several years, this bulletin and its companion on the vaigat formation would not have been the same. we are grateful to the reviewers richard wilson and the late henry emeleus for their helpful and constructive reviews, and to the scientific bulletin editor adam garde for his careful and competent handling of the manuscript. we are highly indebted to our late friend and colleague finn ulff-møller who took part in the field work during several years and who later, during work for the mining company greenex, collected invaluable information in areas not visited by us. special thanks go to the ulff-møller family who handed all his field notes, photographs and other unpublished material over to us after his death. 231 references árting, u.e. 2004: a petrological study of basic 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(eds): volcanism and tectonism in the columbia river flood-basalt province. geological society of america, special paper 239, 131–142. sato, m. 1978: oxygen fugacity and the role of gas-forming elements. geophysical research letters 5, 447–449. saunders, a.d., fitton, j.g., kerr, a.c., norry, m.j. & kent, r.w. 1997: the north atlantic igneous province. in: mahoney, j.j. & coffin, m.l. (eds): large igneous provinces. geophysical monograph 100, 45–93. washington, d.c.: american geophysical union. schiener, e.j. 1975: sedimentological notes on sandstones from nûgssuaq, central west greenland, rapport grønlands geologiske undersøgelse 69, 35–44. schmidt, a.g., riisager, p., abrahamsen, n., riisager, j., pedersen, a.k. & van der voo, r. 2005: palaeomagnetism of the eocene talerua member lavas on hareøen, west greenland. bulletin of the geological society of denmark 52, 27–38. schmincke, h.-u. 1967: fused tuff and peperites in south-central washington. bulletin of the geological society of america 78, 319–330. sjögren, h.j. 1916: om ovifakjärnet och andra telluriska basaltjärn. kungliga svenska vetenskapsakademiens årsbok för år 1916, 255–290. skaarup, n. 2002: evidence for continental crust in the offshore palaeogene volcanic province, central west greenland. geology of greenland survey bulletin 191, 97–102. skaarup, n. & pulvertaft, t.c.r. 2007: aspects of the structure on the coast of the west greenland volcanic province revealed in seismic data. bulletin of the geological society of denmark 55, 65–80. solovova, i.p., ryabchikov, i.d., girnis, a.v., pedersen, a. & hansteen, t. 2002: reduced magmatic fluids in basalt from the island of disko, central west greenland. chemical geology 183, 365–371. steenstrup, k.j.v. 1874: om de kulförende dannelser på öen disko, hareöen og syd-siden af nûgssuak’s halvöen i nord-grönland. videnskabelige meddelelser fra den naturhistoriske forening i kjöbenhavn for aaret 1874, 75–112 + 3 plates. 236236 steenstrup, k.j.v. 1875: om de nordenskiöldske jærnmasser og om forekomsten af gedigent jærn i basalt. videnskabelige meddelelser fra den naturhistoriske forening i kjöbenhavn for aaret 1875, 284–306 + 2 plates. steenstrup, k.j.v. 1883: bidrag til kjendskab til de geognostiske og geographiske forhold i en del af nord-grønland. meddelelser om grønland 4(5), 173–242. steenstrup, k.j.v. 1900: beretning om en undersøgelsesrejse til øen disko i sommeren 1898. meddelelser om grønland 24, 249–306 + plates. storey, m., duncan, r.a., pedersen, a.k., larsen, l.m. & larsen, h.c. 1998: 40ar/39ar geochronology of the west greenland tertiary volcanic province. earth and planetary science letters 160, 569–586. törnebohm, a.e. 1878: über die eisenführenden gesteine von ovifaq und assuk in grönland. bihang till kungliga svenska vetenskapsakademiens handlingar 4(10), 1–22. ulff-møller, f. 1975: high temperature pyrrhotite and telluric iron mineralisations in western disko, central west greenland. rapport grønlands geologiske undersøgelse 75, 51–53. ulff-møller, f. 1977: native iron bearing intrusions of the hammers dal complex, north-west disko. rapport grønlands geologiske undersøgelse 81, 15–33. ulff-møller, f. 1979: new investigations of tertiary lavas and dykes in the area around disko fjord, south disko, central west greenland. rapport grønlands geologiske undersøgelse 95, 30–34. ulff-møller, f. 1983: tellurisk jern fra subvulkanske intrusioner på disko, grønland. en petrografisk og kemisk undersøgelse. unpublished phd thesis, university of copenhagen, 124 pp. + appendices. ulff-møller, f. 1985: solidification history of the kitdlît lens: immiscible metal and sulphide liquids from a basaltic dyke on disko, central west greenland. journal of petrology 26, 64–91. ulff-møller, f. 1986: støbejernsblok på 10 tons fundet på disko. forskning i grønland/tusaat 2/86, 20–25. ulff-møller, f. 1989: exsolution of metallic pb liquid in a magmatic sulphide–metal lens from disko, central west greenland. neues jahrbuch für mineralogie abhandlungen 160, 193–206. ulff-møller, f. 1990: formation of native iron in sediment-contaminated magma. i. a case study of the hanekammen complex on disko island, west greenland. geochimica et cosmochimica acta 54, 57–70. ulff-møller, f. 1991: magmatic platinum-nickel occurrences in the tertiary west greenland basalt province: prospecting by greenex a/s in 1985–1988. open file series grønlands geologiske undersøgelse 91/1, 37 pp. upton, b.g.j. 1988: history of tertiary igneous activity in the n atlantic borderlands. in: morton, a.c. & parson, l.m. (eds): early tertiary volcanism and the opening of the ne atlantic. geological society (london) special publication 39, 429–453. vandenberghe, n., hilgen, f.j. & speijer, r.p. 2012: the paleogene period. in: gradstein, f.m. et al. (eds): the geologic time scale 2012, 855–921. amsterdam: elsevier, http://dx.doi.org/10.1016/ b978-0-444-59425-9.00028-7 white, j.d.l. & houghton, b.f. 2006: primary volcaniclastic rocks. geology 34, 677–680. geological map sheets and sections excepting rosenkrantz et al. (1974, 1976), pulvertaft (1987) and pedersen et al. (2008) all these are also included in the reference list because they are referred to in the main text. pedersen, a.k. & ulff-møller, f. 1987: geological map of greenland, 1:100  000, mellemfjord 69 v.1 nord. copenhagen: geological survey of greenland. pedersen, a.k., larsen, l.m. & dueholm, k.s. 1993: geological section along the south coast of nuussuaq, central west greenland. 1:20 000 coloured geological sheet. copenhagen: geological survey of greenland. pedersen, a.k., ulff-møller, f., larsen, l.m., pedersen, g.k. & dueholm, k.s. 2000: geological map of greenland, 1:100 000, uiffaq 69 v.1 syd. copenhagen: geological survey of denmark and greenland. pedersen, a.k., larsen, l.m., ulff-møller, f., pedersen, g.k. & dueholm, k.s. 2001: geological map of greenland, 1:100 000, pingu 69 v.2 nord. copenhagen: geological survey of denmark and greenland. pedersen, a.k., larsen, l.m. & dueholm, k.s. 2002: geological section along the north side of the aaffarsuaq valley and central nuussuaq, central west greenland. 1:20 000 coloured geological sheet. copenhagen: geological survey of denmark and greenland. pedersen, a.k., larsen, l.m., pedersen, g.k., heinesen, m.v. & dueholm, k.s. 2003: geological section along the south and south-west coast of disko, central west greenland. 1:20 000 coloured geological sheet. copenhagen: geological survey of denmark and greenland. pedersen, a.k., larsen, l.m., pedersen, g.k. & dueholm, k.s. 2005: geological section across north central disko from nordfjord to pingu, central west greenland. 1:20 000 coloured geological sheet. copenhagen: geological survey of denmark and greenland. pedersen, a.k., larsen, l.m., pedersen, g.k., sønderholm, m., midtgaard, h.h., pulvertaft, t.c.r. & dueholm, k.s. 2006: geological section along the north coast of the nuussuaq peninsula, central west greenland. 1:20 000 coloured geological sheet. copenhagen: geological survey of denmark and greenland. pedersen, a.k., pedersen, g.k., larsen, l.m., pulvertaft, t.c.r., sønderholm, m., & dueholm, k.s. 2007a: geological map of the nuussuaq basin in southern nuussuaq, 1:100  000, special map paatuut, with detailed sections. copenhagen: geological survey of denmark and greenland. pedersen, a.k., pedersen, g.k., larsen, l.m., pulvertaft, t.c.r., sønderholm, m., & dueholm, k.s. 2007b: geological map of the south-east coast of nuussuaq between ataata kuua and saqqaqdalen, central west greenland, 1:50  000, with detailed sections. copenhagen: geological survey of denmark and greenland. http://dx.doi.org/10.1016/b978-0-444-59425-9.00028-7 http://dx.doi.org/10.1016/b978-0-444-59425-9.00028-7 237 pedersen, a.k., larsen, l.m., pedersen, g.k. & dueholm, k.s. 2008: geological map of the area around sikillingi, western nuussuaq, central west greenland, 1:20 000. copenhagen: geological survey of denmark and greenland. pulvertaft, t.c.r. 1987: geological map of greenland, 1:100 000, agpat 70 v.2 nord. copenhagen: geological survey of denmark and greenland. rosenkrantz, a., münther, v. & henderson, g. 1974: geological map of greenland, 1:100 000, agatdal 70 v.1 nord. københavn: grønlands geologiske undersøgelse. rosenkrantz, a., münther, v., henderson, g., pedersen, a.k. & hald, n. 1976: geological map of greenland, 1:100 000, qutdligssat 70 v.1 syd. københavn: grønlands geologiske undersøgelse. 238238 appendix: place names the place names mentioned in the text and listed below are shown on the maps in fig. 1 (regional names) and figs 4, 6 and 103 (local names and localities). aasiaat ............................................................................. fig. 1 aaffarsuaq .................................................................. figs 4, 6 agatdalen ........................................................................ fig. 6 agatfjeldet ....................................................................... fig. 6 akuarut (lyngmarksfjeld) .......................................... fig. 6 akuliarusersuaq ........................................................ figs 4, 6 akulliit (mellemfjord) ............................................ figs 4, 6 akunneq ......................................................................... fig. 6 alanngup qaqqai .......................................................... fig. 6 apostelfjeld (navaranaat) ........................................... fig. 6 aqajaruata qaqqaa .................................................. figs 4, 6 assoq .......................................................................... figs 4, 6 asuk ............................................................................ figs 4, 6 asuutaa ....................................................................... figs 4, 6 ataata kuua ............................................................... figs 4, 6 blæsedalen ...................................................................... fig. 6 blåbærdalen ............................................................... figs 4, 6 brededal ..................................................................... figs 4, 6 charles polaris dal ....................................................... fig. 6 daugaard-jensen dal ............................................... figs 4, 6 disko bugt ...................................................................... fig. 1 disko gneiss ridge ....................................................... fig. 4 disko fjord (kangerluk) ......................................... figs 4, 6 enok havn (ivisaarqut) ................................................ fig. 6 eqaluit (nordre laksebugt) ................................... figs 4, 6 eqalunnguaqqat qaqqaat ...................................... figs 4, 6 eqi ............................................................................... figs 4, 6 eqip qaqqaa ............................................................. figs 4, 6 fortunebay (killiit) ................................................. figs 4, 6 frederik lange dal .................................................. figs 4, 6 gamle qullissat (qullissaaqqat) ............................... fig. 6 giesecke dal ..................................................... figs 4, 6, 103 giesecke monument (uppalluk) ......................... figs 4, 6 godhavn (qeqertarsuaq) ....................................... figs 4, 6 hammer dal .................................................... figs 4, 6, 103 hanekammen .................................................. figs 4, 6, 103 hareøen .................................................................figs 1, 4, 6 ikorfarsuit .................................................................. figs 4, 6 ikorfat ......................................................................... figs 4, 6 illukasik ........................................................................... fig. 6 illukunnguaq .................................................................. fig. 6 illuluarsuit qaqqaa .............................................. fig. 6, 103 innerit .............................................................................. fig. 1 inngigissoq ................................................................ figs 4, 6 innarsuaq (skarvefjeld) .......................................... figs 4, 6 inussuk ............................................................................ fig. 6 ippik ................................................................................. fig. 6 itilli (fault) ................................................................. figs 1, 4 ivisaarqut (enok havn) ............................................... fig. 6 ivissussat qaqqaat .................................................... figs 4, 6 jamma .............................................................................. fig. 6 jernpynten ...................................................................... fig. 6 kangerluk .................................................................. figs 4, 6 kangersooq (nordfjord) ........................................ figs 4, 6 keglen .............................................................................. fig. 6 killerpaat qaqqarsuat .................................................. fig. 6 killiit (fortunebay) ................................................. figs 4, 6 kingittup qaqqaa ......................................................... fig. 6 kingittuusaq .........................................................figs 6, 103 kuannersuit kuussuat .................................................. fig. 6 kuannersuit sulluat ...................................................... fig. 6 kuannit ........................................................................... fig. 6 kuugannguaq ............................................................ figs 4, 6 kvandalen .................................................................. figs 4, 6 laksedalen ................................................................. figs 4, 6 luciefjeld ........................................................................ fig. 6 lyngmarksfjeld (akuarut) .......................................... fig. 6 makittarissagaq ............................................................. fig. 6 maniillat .......................................................................... fig. 6 marraat qaqqaat ...................................................... figs 4, 6 mellemfjord (akullit) ............................................. figs 4, 6 morten porsild dal .............................................. fig. 6, 174 naqerloq ......................................................................... fig. 6 narsap qaqqaa ......................................................... figs 4, 6 navaranaat (apostelfjeld) ........................................... fig. 6 niaqussat ................................................................... figs 4, 6 niiortuut .................................................................... figs 4, 6 niuluut ........................................................................ figs 4, 6 nordfjord (kangersooq) ......................................... figs 4, 6 nordre laksebugt (eqaluit) ................................... figs 4, 6 nunavik ...................................................................... figs 4, 6 nuuk kangilleq ............................................................. fig. 6 239 orlingasoq ................................................................. figs 4, 6 orpiit qaqqaat ......................................................... figs 4, 6 paatuut ....................................................................... figs 4, 6 perlertut qaqqaat ..................................................... figs 4, 6 pingu ........................................................................... figs 4, 6 point 440 m ...........................................................figs 6, 103 point 500 m ...........................................................figs 6, 103 point 600 m ............................................................... fig. 103 point 780 m .................................................................... fig. 6 point 882 m ...........................................................figs 6, 103 point 975 m .................................................................... fig. 6 point 1014 m ............................................................. figs 4,6 point 1025 m ................................................................. fig. 6 point 1070 m ........................................................figs 6, 103 point 1109 m ................................................................. fig. 6 point 1123 m ............................................................ figs 4, 6 point 1132 m ........................................................figs 6, 103 point 1137 m ........................................................figs 6, 103 point 1266 m ................................................................. fig. 6 point 1300 m ............................................................ figs 4, 6 point 1380 m ........................................................figs 6, 103 point 1440 m ................................................................. fig. 6 point 1460 m ........................................................figs 6, 103 point 1510 m .........................................................figs 6, 103 point 1530 m ............................................................ figs 4, 6 point 1560 m ............................................................ figs 4, 6 point 1578 m ............................................................ figs 4, 6 point 1590 m ........................................................figs 6, 103 point 1610 m ........................................................figs 6, 103 point 1640 m ............................................................ figs 4, 6 point 1650 m ................................................................. fig. 6 point 1722 m ................................................................. fig. 6 point 1760 m ................................................................. fig. 6 point 1888 m .............................................................fig. 4, 6 point 2000 m ................................................................. fig. 6 point 2010 m ................................................................. fig. 6 point 2080 m .............................................................fig. 4, 6 puiattussuaq ................................................................... fig. 6 pyramiden ................................................................. figs 4, 6 qasigissat (sælbugten) ............................................ figs 4, 6 qasigissat kuussuat (vesterdalen) ........................ figs 4, 6 qeqertarsuaq (godhavn) ....................................... figs 4, 6 qinngusaq ................................................................. figs 4, 6 qullissaaqqat (gamle qullissat) ............................... fig. 6 qullissat ..................................................................... figs 4, 6 qunnilik ......................................................................... fig. 6 rink dal ............................................................ figs 4, 6, 103 rødeelv ........................................................................... fig. 6 sapernuvik ................................................................. figs 4, 6 saqqaq ........................................................................ figs 4, 6 saqqarliit ilorliit ....................................................... figs 4, 6 saqqarliit silarliit .......................................................... fig. 6 sedimentkløften ...................................................figs 6, 103 sermersuaq ..................................................................... fig. 6 skarvefjeld (innarsuaq) .......................................... figs 4, 6 skorstensfjeld ............................................................ figs 4, 6 siniffik ............................................................................. fig. 6 sorte hak ................................................................... figs 4, 6 sortebærdalen ........................................................... figs 4, 6 steenstrup dal ......................................................figs 6, 103 stordal ........................................................................ figs 4, 6 svartenhuk halvø ......................................................... fig. 1 tartunaq ..................................................................... figs 4, 6 tini .................................................................................. fig. 6 tuapassuit ....................................................................... fig. 6 tuapaat ............................................................................ fig. 6 tuapaat qaqqaat ...................................................... figs 4, 6 tunup qaqqaa .......................................................... figs 4, 6 ubekendt ejland ........................................................... fig. 1 uiffaq ......................................................................... figs 4, 6 ukaleqartarfik ............................................................... fig. 6 umiartorfiup qaqqaa .................................................. fig. 6 umiusat ...................................................................... figs 4, 6 uppalluk (giesecke monument) .......................... figs 4, 6 uunartuarsuk ................................................................. fig. 6 vaigat ......................................................................figs 1, 4, 6 vesterdalen ................................................................ figs 4, 6 de nationale geologiske undersøgelser for danmark og grønland (geus) geological survey of denmark and greenland g eological survey of d enm ark and g reenland b ulletin 40 • 2018 volcanic rocks of the paleocene m aligât form ation, w est g reenland øster voldgade 10 dk-1350 copenhagen k denmark lithostratigraphy, geology and geochemistry of the volcanic rocks of the maligât formation and associated intrusions on disko and nuussuaq, paleocene of west greenland the upper cretaceous–tertiary nuussuaq basin in west greenland contains a many kilometres thick succession of siliciclastic sediments and overlying volcanic rocks. the first studies in the early 19th century were centred on the coal and fossils in the sediments and the minerals in the volcanic rocks, including famous occurrences of native iron. the present focus of interest includes modern stratigraphic and volcanological studies to decipher the basin evolution and support hydrocarbon and mineral exploration. this bulletin presents the lithostratigraphy, geology and geochemistry of the paleocene volcanic rocks of the maligât formation and its related intrusions on disko and the nuussuaq peninsula; it concludes with a detailed discussion of the effects of crustal contamination processes. the maligât formation is up to 2000 m thick and comprises four formally defined members and 15 chemically defined informal units. it is mainly composed of basalt lavas but also includes basaltic andesite, andesite and dacite flows and rhyolite tuffs. the silicic rocks and intrusions were produced by contamination in high-level magma chambers and commonly contain native iron. the comprehensive descriptions and analyses of each member and unit represent a synthesis of many years of work and are intended to serve as a guide for future studies, including exploration for mineral deposits associated with some units of the formation. a companion bulletin (volume 39) on the volcanic rocks of the picritic vaigat formation that underlies the maligât formation was published in 2017. lithostratigraphy lithostratigraphy, geology and geochemistry of the volcanic rocks of the maligât formation and associated intrusions on disko and nuussuaq, paleocene of west greenland asger ken pedersen, lotte melchior larsen & gunver krarup pedersen g e o l o g i c a l s u r v e y o f d e n m a r k a n d g r e e n l a n d b u l l e t i n 4 0 c o p e n h a g e n • 2 0 1 8 printed 2018 issn 1604-8156 isbn 978-87-7871-498-5 geus is a research and advisory institution in the danish ministry of energy, utilities and climate d e n m a r k geocenter denmark is a formalised cooperation between geological survey of denmark and greenland (geus), department of geoscience at aarhus university and the geological museum and department of geosciences and natural resource management at the university of copenhagen. 9 788778 714985 forside_web bulletin40_240s_weblinks bagside_web geological survey of denmark and greenland bulletin 35, 2016, 55-58 55© 2016 geus. geological survey of denmark and greenland bulletin 35, 55–58. open access: www.geus.dk/publications/bull new zircon u-pb and hf isotopic constraints on the crustal evolution of the skjoldungen region, south-east greenland thomas f. kokfelt, tomas næraa, kristine thrane and leon bagas north atlantic craton north atlantic craton skjoldungen 500 km quaternary skjoldungen alkaline province basement rocks gabbro/diorite granite/diorite agmatitic orthogneiss orthogneiss granite/syenite mafic granulite singertât complex ultramafic rock palaeoproterozoic appinite suite syenitic gneiss 41°w 42°w 43°w in lan d ice 6 3 °3 0 'n 6 3 °n 6 2 °3 0 'n 25 km25 km mogens heinesen fjord bernsto rff is fjord skjoldungenskjoldungen fig. 1. geological map of the skjoldungen region, south-east greenland, showing locations of samples (blue circles) selected for zircon u-pb geochronology and zircon hf isotope analysis. we report new zircon u-pb and hf isotopic data from the skjoldungen region between c. 62°30´ and 63°40´n in southeast greenland. the work was carried out under the southeast greenland mineral endowment task (segment); a joint project between the geological survey of denmark and greenland (geus) and the ministry of mineral resources (mmr) in greenland to assess the mineral endowment and update the geological knowledge of the region using modern petrological, geochemical and geochronological tools. this paper presents new zircon u-pb and hf isotopic data from a range of different archaean rocks in the skjoldungen region, which greatly improve the understanding of the history of crustal growth. regional geology the skjoldungen region in south-east greenland as defined here covers the ice-free area between mogens heinesen fjord in the south and bernstorff isfjord in the north (fig. 1). the region exposes a midto lower-crustal section of the archaean north atlantic craton (kolb et al. 2013 and references therein). the dominant rock type is granodiorite with lesser amounts of monzogranite and rare tonalite, commonly with nebulitic to agmatitic textures, and with abundant mafic and ultramafic inclusions. the northern part of the region includes grey orthogneiss that is interleaved with the agmatitic gneiss. the southern part is structurally highly complex containing abundant migmatitic rocks, generally recording lower crustal conditions. remnants of older supracrustal rocks forming kilometre-sized lensoidal belts of mafic granulite, ultramafic rocks and paragneiss are found in the northern and southernmost parts of the region (fig. 1). these rocks are commonly invaded by felsic partial melts, dismembering their border zones into smaller units that also account for the agmatitic texture of the surrounding gneiss. kolb et al. (2013) and bagas et al. (2013) proposed that the region was deformed during both the > c. 2800 ma timmiarmiut and the c. 2790–2700 ma skjoldungen orogenies. the timmiarmiut orogeny is only weakly defined based on deformation events in the southern part of the area that are overprinted by the skjoldungen orogeny. skjoldungen island and its surroundings reached granulite facies peak conditions at c. 2760–2740 ma. subsequent orogenic collapse characterised by fast exhumation rates (berger et al. 5656         2500 2700 2900 3100 3300 3500 3700 3900 inherited and gneiss protolith age granitoid intrusion and pegmatite/aplite   all data (109 samples, 178 ages)  2 7 1 0 2 6 9 5 2 7 4 0 2 7 5 0 2 9 1 5 3 0 1 5 3 8 8 52 8 0 0 3 2 3 0 3 1 6 0 3 1 0 0 0 0.01 0.02 a 62.2 62.6 63.0 63.4 63.8 l at it u d e °n b r el at iv e p ro b ab ili ty 207pb/206pb age (ma) skjoldungen alkaline province ‘timmiarmiut orogeny’ (accretionary orogen?) inherited and gneiss protolith age granitoid intrusion and pegmatite/aplite   fig. 2. a: 207pb/206pb age distribution diagram for the skjoldungen region, based on 109 samples that provide 178 distinct ages. the black dashed curve indicates the relative probability of all age data. b: 207pb/206pb ages plotted against geographical latitude. the rare, detected oldest age components are restricted to the southern part of the region, whereas the dominant ≤3100 ma age components are equally distributed. 2014) resulted in extensive crustal remelting and continued emplacement of mildly alkaline plutons forming the skjoldungen alkaline province, which constitutes a globally rare occurrence of archaean alkaline magmatism (blichert-toft et al. 1996). the province comprises a large number of mafic and ultramafic to differentiated intrusions within a large area of c. 2400 km2 centered at the wnw-trending skjoldungen island (blichert-toft et al. 1995). the magmatism comprises a slightly alkaline, c. 2750–2690 ma, stage followed by a highly alkaline, c. 2680-2664 ma singertât stage, which includes nephelinitic and carbonatitic rocks (nielsen & rosing 1990; nutman & rosing 1994; kolb et al. 2013). samples and methods a total of 109 samples of orthogneiss, migmatite, granitic rocks, pegmatite and aplite were collected for u-pb zircon geochronology. as seen in fig. 1 the geographical sample distribution is uneven, with most samples collected in the northern part of the region. similarly, coastal areas are more densely sampled than remote and less accessible areas near the ice sheet. most u-pb data were acquired using laser ablation-single collector-magnetic sector field-inductively coupled plasmamass spectrometry (la-sf-icp-ms) at geus, employing a thermo finnigan element2 mass spectrometer coupled to a new wave research up213 frequency-quintupled solid state nd:yag laser system. the analytical procedures followed frei & gerdes (2009) and dziggel et al. (2014). a few samples were analysed at the nordsim facility at the natural history museum in stockholm using standard procedures in whitehouse et al. (1999), and at the shrimp facility at curtin university using standard procedures in compston et al. (1984). different textural domains of the zircon grains such as cores and rims were documented using scanning electron microscopy and targeted during analysis. hf isotope data by la-icpms were obtained on a subset of samples, targeting the same textural domains as were analysed for u-pb. the hf isotope data were analysed at the university of frankfurt following the methods in gerdes & zeh (2006). results: u-pb data more than 7000 analytical spots on zircon grains from the 109 samples resulted in 178 interpreted u-pb ages, reflecting that some samples contained more than one age population. intrusion ages were calculated from analyses that were 90–110% concordant, or from upper concordia intercepts. single spot ages interpreted as inherited should be treated as minimum ages due to potential ancient pbloss. the u-pb age data are summarised in a density distribution diagram (fig. 2a) where the data define two age groups: a group (shown in red) of granitic veins, sheets and plutons, pegmatite and aplite dykes contemporaneous with the skjoldungen orogeny, and a group (shown in blue) of pre-skjoldungen ages related to both grey orthogneiss samples and inherited zircon grains. grains that are older than the main intrusive age are interpreted as inherited. in the southern area such grains are thought to represent older, reworked gneiss protoliths. in the skjoldungen alkaline province inherited zircons are likely to reflect crustal contamination of juvenile melts. overall, the dataset spans about 1.3 ga of earth history from c. 3880 to 2600 ma, but with the overwhelming majority of ages between c. 3250 and 2690 ma. a major age peak between c. 2750 and 2690 ma reflects that about one third of the analysed samples were collected from the skjoldungen alkaline province. at closer inspection, this relatively narrow age range may be divided into several distinct age peaks at c. 2750, 2740, 2710 and 2695 ma (fig. 2a). 57 mesoto eoarchaean source mainly mesoarchaean sourceskjoldungen alkaline province g ra n it o id a n d g n ei ss ic b as em en t b 25 km mafic crust depleted mantle model age c. 4000 ma -12 -8 -4 0 4 2700 2900 3100 3300 εh f (t ) depleted mantle chur 207pb/206pb age (ma)a fig. 3. geographical distribution of zircon analytical data. a: zircon εhf plotted against 207pb/206pb ages. samples collected north and south of 63° 10´ shown in green and brown. the compositional range of the skjoldungen alkaline province (blue field; næraa et al. 2015) corresponds to a nearchondritic signature and overlaps with the data from the northern part of the region. chur: chondritic uniform reservoir. b: division of the skjoldungen region into geographical areas based on zircon hf isotopic compositions. the northern part is derived from a mainly mesoarchaean source (2950–3100 ma), and the southern part from a mainly mesoand eoarchaean source (≥3200 ma). the zircon ages are plotted against geographical latitude in fig. 2b. the diagram displays a systematic difference in the age distribution of inherited zircons north and south of 63°10´n, despite the above-mentioned uneven sampling density. the southern and northern areas share a significant inherited population with ages spanning from c. 3100 to 2800 ma, but whereas the northern area has few zircon grains older than c. 3100 ma, the southern area contains several inherited grains with ages between c. 3880 and 3100 ma. the diagram also arguably reveals a broad trend of northwards decreasing maximum ages. a similar, albeit less pronounced and more local, trend is seen for the younger intrusions of skjoldungen alkaline province, where the oldest ages of c. 2750 ma are found in the south (syenitic gneiss in the skirner bjerge and kassertoq areas), and the ages decrease down to c. 2720 ma in the north (figs 1, 2b). these age trends reflect regional differences in the crust at depth and have implications for the interpretation of the regional geodynamic evolution. results: hf isotope data combined u-pb and hf zircon isotope data have been obtained from 29 samples (fig. 3). the data in fig. 3a are displayed in fig. 3b by assigning individual samples to either the northern or southern area by combining intrusive and inherited ages for each area for simplicity. the northern area is characterised by relatively higher initial εhf values and is interpreted as being more juvenile, with a depleted mantle model age of c. 3100 ma. the southern area has somewhat lower initial εhf values and has older model ages of c. 4000–3300 ma (assuming a linear depleted mantle model from εhf = 0 at 4500 ma to +17 at present), and is interpreted as less juvenile. although both areas also share an age distribution between 3100 and 2700 ma, the hf isotope data indicate that the melts in the northern and southern areas were sourced from different regions, with the northern area tapping a relatively more juvenile and younger crustal reservoir than the southern area. discussion and summary the comprehensive new u-pb and hf isotope dataset from the skjoldungen region sheds new light on the tectono-magmatic evolution of the north atlantic craton. as illustrated in figs 2b and 3, the skjoldungen region can be divided into two crustal terranes with distinct hf isotopic signatures, albeit with partially overlapping age profiles. both terranes include u-pb zircon ages between 3100 and 2700 ma, but the southern terrane also contains a noticeable component of inherited grains between c. 3880 and 3200 ma (figs 2, 3). traces of such old crustal remnants have rarely been found in south-east greenland but have been recorded from local stream sediments (thrane & keulen 2015). although the basement throughout the skjoldungen region records a similar age distribution from c. 3100 ma onwards, the zircon hf isotopic data define two distinct crustal source terranes, where the zircon hf isotope data in these rocks display higher initial εhf values in the north than in the south (fig. 3). the few inherited zircon ages from c. 3880 to 3100 ma leave the geodynamic setting responsible for the earliest magmatic activity unclear. still, the u-pb-hf isotope data suggest long-lived growth of a terrane or crustal block that includes very old (>4000 ma) crustal source components. the subsequent period from c. 3100 to 2800 ma records semi-continuous magmatic activity throughout the entire skjoldungen region with contemporaneous development of rocks of the northern and southern terranes and diminishing juvenile input over time, as reflected by the steady decrease in hf isotope compositions (fig. 3a). these signatures could reflect that the northern terrane formed in 5858 authors’ addresses t.f.k. & k.t., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: tfk@geus.dk t.n., department of geology, lund university, sölvegatan 12, se-223 62 lund, sweden. l.b., school of earth and environment, university of western australia, 35 stirling highway, crawley wa 6009, australia. an accretionary orogeny that was building out from south to north. in this model, south-verging subduction along the northern boundary of the southern terrane would have proceeded for several hundreds of millions of years with progressive accretion of arc systems to form the northern juvenile terrane. alternatively, the northern terrane with its distinct and younger hf source signature could have developed as a separate entity prior to the skjoldungen orogeny, in which case the broad overlap in ages between the southern and northern terranes would be coincidental. the prolonged period of accretion-related magmatism diminished at c. 2800 ma, where the prelude to the skjoldungen orogeny represents a shift in tectonic setting to that of a continent–continent collision orogeny (bagas et al. 2013; kolb et al. 2013). the orogeny was associated with renewed, intensified magmatism and the development of the skjoldungen alkaline province (blichert-toft et al. 1995). the new zircon data from the main magmatic stage of the skjoldungen alkaline province are consistent with formation in a protracted time period from c. 2750 to 2690 ma, overlapping with the late stage of the skjoldungen orogeny. whole rock hf isotope data from mafic intrusions in the province plot with near-primitive mantle signatures (εhf from 0 to –2) and overlap with the hf zircon isotope composition of the northern area (næraa et al. 2015). the slightly negative hf isotope values are consistent with derivation from a slightly enriched mantle source, probably coupled with an enrichment in incompatible elements arising from the preceding subduction processes, and by assimilation of crustal material. the new age data from the province also suggest distinct magmatic events at c. 2750, 2740, 2710 and 2695 ma. the reason for such episodic magmatism is currently not understood but might be tectonically controlled. the new u-pb-hf isotopic data dramatically widen the insight into the crustal development of south-east greenland. the new data indicate that the study region consists of two different domains, a southern one containing source components older than c. 3200 ma (up to c. 4000 ma), and a northern one entirely younger than c. 3100 ma. these systematics might either be explained by the existence of two distinct crustal terranes that were amalgamated at c. 2800 ma, or as one continuous terrane that was built out by progressive accretion from south to north between c. 3100 and 2800 ma. references bagas, l., næraa, t., kolb, j., reno, b.l. & fiorentini, m.l. 2013: partial melting of the archaean thrym complex of southeastern greenland. lithos 160–161, 164–182. berger, a., kokfelt, t.f. & kolb, j. 2014: exhumation rates in the archean from pressure–time paths: example from the skjoldungen orogen (se greenland). precambrian research 255, 774–790. blichert-toft, j., rosing, m.t., lesher, c.e. & chauvel, c. 1995: geochemical constraints on the origin of the late archaean skjoldungen alkaline igneous province, se greenland. journal of petrology 36, 515–561. blichert-toft, j., arndt, n.t. & ludden, j.n. 1996: precambrian alkaline magmatism. lithos 37, 97–111. compston w., williams i. & meyer c. 1984: u-pb geochronology of zircons from lunar breccia 73217 using a sensitive high mass-resolution ion microprobe. journal of geophysical research: solid earth (1978–2012) 89, b525–b534. dziggel, a., diener, j.f.a., kolb, j. & kokfelt, t.f. 2014: metamorphic record of accretionary processes during the neoarchaean: the nuuk region, southern west greenland. precambrian research 242, 22–38. frei, d. & gerdes, a. 2009: precise and accurate in situ u–pb dating of zircon with high sample throughput by automated la-sf-icp-ms. chemical geology 261, 261–270. gerdes, a. & zeh, a. 2006: combined u–pb and hf isotope la-(mc-) icp-ms analyses of detrital zircons: comparison with shrimp and new constraints for the provenance and age of an armorican metasediment in central germany. earth and planetary science letters 249, 47–61. kolb, j., thrane, k. & bagas, l. 2013: field relationship of high-grade neoto mesoarchaean rocks of south-east greenland: tectonometamorphic and magmatic evolution. gondwana research 23, 471–492. nielsen, t.f.d. & rosing, m.t. 1990: the archaean skjoldungen alkaline province, south-east greenland. rapport grønlands geologiske undersøgelse 148, 93–100. nutman, a & rosing, m.t. 1994: shrimp u-pb zircon geochronology of the archaean ruinnæsset syenite, skjoldungen alkaline province, south-east greenland. geochimica et cosmochimica acta 58, 3515–3518. næraa, t., bagas, l., tusch, j., kokfelt, t.f. & münker, c. 2015: pro and retrograde igneous activity during the neoarchaean skjoldungen orogeny in the se greenland. goldschmidt abstracts 2248. thrane, k. & keulen, n. 2015: provenance of sediments in the faroe– shetland basin: characterisation of possible source components in southeast greenland. in: 5th faroe islands exploration conference: proceedings of the 4th conference, 7–25. annales societatis scientiarum færoensis supplementum lxiv. whitehouse, m.j., kamber, b.s. & moorbath, s. 1999: age significance of u–th–pb zircon data from early archean rocks of west greenland – a reassessment based on combined ion-microprobe and imaging studies. chemical geology 160, 210–224. geological survey of denmark and greenland bulletin 13, 2007, 17-20 17 intense investigations of deep aquifers in jylland, western denmark, during the last seven years have resulted in de tailed mapping of miocene sand-rich deposits laid down in fluvial channels, delta lobes, shoreface and spit complexes (fig. 1; rasmussen 2004). detailed sedimentological and paly no logical studies of outcrops and cores, and interpretation of high-resolution seismic data, have resulted in a well-founded sequence-stratigraphic and lithostratigraphic scheme (fig. 1) suitable for prediction of the distribution of sand. the miocene succession onshore denmark is divided into three sand-rich deltaic units: the ribe and bastrup sands and the odderup formation (fig. 2). prodeltaic clayey deposits of the vejle fjord and arnum formations interfinger with the sand-rich deposits. most of the middle and upper mio cene in denmark is composed of clayey sediments referred to the hodde and gram formations (fig. 2). this paper presents examples of seismic reflection patterns that have proved to correlate with sand-rich deposits from lower miocene deltaic deposits and that could be applied in future exploration for aquifers and as analogues for oiland gas-bearing sands in wave-dominated deltas. geology during the early miocene, the eastern north sea basin was filled by siliciclastic sediments sourced from the fenno scandian shield. the sediment supply was high due to tectonic uplift of the fennoscandian shield (ziegler 1990; rasmussen 2004). the north sea was located in the highlati tude belt of westerly winds, which resulted in a long fetch, and the tidal range is interpreted to be microto meso-tidal. regressions and transgressions during the early miocene were strongly controlled by eustatic sea-level changes (friis et al. 1998; rasmussen 2004; rasmussen & dybkjær 2005). during the early miocene, two phases of shoreline progradation occurred. sand deposited adjacent to the delta mouth or in association with topographic highs was immediately redistributed and deposited either as spit complexes or as barrier islands in the down-drift areas of delta lobes. these sandrich successions are commonly around 20 m thick; however, delta lobes prograding into topographic lows, i.e. deep water, are characterised by up to 70 m thick successions of clean sand. the delta front sediments were deposited either as mass-flow sediments or current-derived deposits. during sealevel fall, incision of the delta plain took place. these incised valleys were successively filled with thick, sand-rich fluvial deposits during the succeeding sea-level rise. prediction of reservoir sand in miocene deltaic deposits in denmark based on high-resolution seismic data erik s. rasmussen, thomas vangkilde-pedersen and peter scharling © geus, 2007. geological survey of denmark and greenland bulletin 13, 17–20. available at: www.geus.dk/publications/bull fig. 1. map of jylland showing distribution of lower miocene environments (from rasmussen 2004). insert map shows position of seismic sections and boreholes used in this study. seismic data acquisition mapping of aquifers in denmark has previously been dominated by electric and electromagnetic methods, as the high cost of conventional, shallow, onshore reflection seismic surveys was a factor that limited its use. recently, however, the technique of landstreamer high-resolution seismic data has provided considerable savings of manpower and increased productivity compared to using traditionally planted geophones and cable lay-outs. the mapping of deeper units is also possible now. the landstreamer technique also facilitates short geophone spacing and differential spacing of geophones along the spread without increasing timeor manpower consumption. the use of landstreamers for acquisition of shallow seismic data has increased throughout the world in recent years. the landstreamers are commonly used together with relatively weak sources such as a pipegun or sledgehammer (e.g. van der veen & green 1998; van der veen et al. 2001) resulting in a relatively limited penetration depth (typically a few hundred metres). since the year 2000, more than 1000 km of highresolution seismic data have been acquired to map deep aquifers in denmark. the acquisition setup used has included high-frequency seismic vibrators (3.5 t and 6.5 t) as the energy source. under normal conditions the landstreamer setup has provided very high data quality with reflections from c. 20–50 m down to more than 1 km with a vertical resolution of 5–10 m. the coverage, especially in the central and western parts of jylland, provides unique opportunities for interpretation and correlation. the design of the seismic landstreamers has developed from a 150 m, 60-channel streamer with 2.5 m spacing used 18 fig. 2. lithostratigraphy of the danish miocene sediments (modified from rasmussen 2004). fig. 3. seismic section with boreholes from the billund area illustrating two prograding deltaic sand-rich units (billund sand and bastrup sand). the grain size of the penetrated succession is indicated by different colours. note that the parallel clinoformal seismic reflection pattern (arrows) always correlates with sand. seismic data courtesy of cowi a/s and rambøll a/s. in 2000 to the current 200–220 m streamers with 96–102 channels and differential geophone spacings of 1.25, 2.5 and 5 m (vangkilde-pedersen et al. 2003, 2006). the differential geo phone spacing along the streamers, with the shortest spacing close to the vibrator, has greatly improved the quality and reso lution of the near-surface data. in the same period, both the vibrator sweeps and processing of the data have also been optimised. during the first couple of years a simple standard processing sequence was applied to the data, but in recent years the processing sequence has been significantly improved. examples of lower miocene reservoir sand delta front sand. thick delta front sands occur in association with progradation into deep water that is normally associated with structurally confined areas. delta lobes deposited during a relative sea-level fall are especially sand-rich. these deposits are characterised by a parallel clinoform reflection pattern (figs 3, 4) in which the dip of the clinoforms range from 7° to 10°. the thickness of sand associated with this reflection pattern has never been recorded as less than 20 m and thicknesses of up to 50 m have been found at billund (fig. 3); the thickness may be more than 70 m within the brande lobe, north of billund. the grain size is commonly medium to coarse sand, but gravel may occur in connection with mass-flow deposits on the delta front or in association with channels and shoreface deposits in the upper part of the delta. delta sand laid down during a sea-level fall is particularly clean and homogenous. fluvial point-bar sand. delta deposits of the lower miocene bastrup sand are often capped by fluvial sediments that have been protected during the succeeding transgression. the fluvial channels are expressed by a concave-up structure filled with a shingled seismic reflection pattern probably representing point-bar deposits (fig. 5). from seismic and borehole data, the point-bar deposits comprise up to 20 m thick, fining-upwards successions composed of coarseto finegrained sand that are commonly capped by coal. incised valley sand. well-defined, large, concave-upward erosional surfaces are found especially in the proximal parts of the deltas. the infills of these features on the seismic lines are often characterised by a transparent seismic reflection pattern (fig. 6). from outcrop and borehole data the valleys are known to be filled typically by coarse-grained sand and gravel that were deposited in braided river systems. the thickness and lateral distribution of the fill vary within the delta complexes. future perspectives the application of seismic data in the search for aquifers in denmark by recognition of different seismic reflection pat19 fig. 4. seismic section from sønder omme. note the close relationship between the parallel clinoformal reflection pattern and the sand as indicated by the arrows. note especially the lower delta where only the toe of the delta front has been penetrated by the stakroge borehole. seismic data courtesy of rambøll a/s. see fig. 3 for legend. fig. 5. seismic section showing shingled, seismic reflection pattern (arrow) within a channel structure. similar structures have been found on 3d seismic data from canada (posamentier 2005) and represent late ral accretion of a point bar. seismic data courtesy of cowi a/s. 20 terns and morphological features, e.g. the geometry of clinoforms, has proved to be useful in the prediction of sand-rich sediments in miocene deposits. sand-rich sediments in front of a delta complex are normally associated with clinoform reflection patterns. a shingled seismic reflection pattern within channels characterises sand-rich, point-bar deposits. distinct erosional features, with a transparent reflection pattern, capping delta foresets commonly indicate fluvial sandrich sediments. a detailed mapping of the miocene delta complexes and the construction of a three-dimensional model of delta lobes will be essential for developing future hydrogeological models. furthermore, the connection between the observed seismic facies and sand-rich environments may also be applied as a tool for prediction of jurassic hydrocarbon reservoir sands in the north sea area. acknowledgements the carlsberg foundation and the counties of vejle, ringkøbing and ribe are thanked for financial support of the study of the miocene succession in denmark. references friis, h., mikkelsen, j. & sandersen, p. 1998: depositional environment of the vejle fjord formation of the upper oligocene – lower miocene of denmark: a back island/barrier-protected depositional complex. sedimentary geology 17, 221–244. posamentier, h.w. 2005: application of 3d seismic visualization techniques for seismic stratigraphy, seismic geomorphology and depositional systems analysis: examples from fluvial to deep-marine depo sitional environments. in: doré, a.g. & vining, b.a. (eds): petroleum geology: north-west europe and global perspectives – proceedings of the 6th petroleum geology conference, 1563–1576. london: geo logical society. rasmussen, e.s. 2004: stratigraphy and depositional evolution of the uppermost oligocene – miocene succession in western denmark. bulletin of the geological society of denmark 51, 89–109. rasmussen, e.s. & dybkjær, k. 2005: sequence stratigraphy of the upper oligocene – lower miocene of eastern jylland, denmark: role of structural relief and variable sediment supply in controlling sequence development. sedimentology 52, 25–63. van der veen, m. & green, a.g. 1998: landstreamer for shallow seismic data acquisition: valuation of gimbal-mounted geophones. geophy sics 63, 1408–1413. van der veen, m., spitzer, r., green, a.g. & wild, p. 2001: design and application of a towed landstreamer for cost-effective 2d and pseudo3d shallow seismic data acquisition. geophysics 66, 482–500. vangkilde-pedersen, t., skjellerup, p., ringgaard, j. & jensen, j.f. 2003: pulled array seismic (pas) – a new method for shallow reflection seismic data acquisition. 65th eage conference & exhibition, stavanger, norway, 2–5 june 2003. extended abstracts, 201 only. vangkilde-pedersen, t., dahl, j.f. & ringgaard, j. 2006: five years of experience with landstreamer vibroseis and comparison with conventional seismic data acquisition. proceedings of the 19th annual sageep symposium on the application of geophysics to engineering and environmental problems, seattle, usa, 1086–1093. ziegler, p. 1990: geological atlas of western and central europe, 2nd edition, 239 pp, mijdrecht: shell international petroleum. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: esr@geus.dk fig. 6. seismic section showing 2 km wide and 20 m deep erosional features on top of clinoforms interpreted as a fluvial valley fill. these features are often filled with coarse-grained sand or gravel deposits. seismic data courtesy of cowi a/s. borehole seismic studies of a volcanic succession from the lopra-1/1a borehole in the faroe islands, northern north atlantic 23 borehole seismic studies of a volcanic succession from the lopra-1/1a borehole in the faroe islands, northern north atlantic philip christie, ian gollifer and david cowper extruded basalt flows overlying sedimentary sequences present a challenge to hydrocarbon exploration using reflection seismic techniques. the lopra-1/1a re-entry well on the faroese island of suðuroy allowed us to study the seismic characteristics of a thick sequence of basalt flows from well logs and borehole seismic recordings. data acquired during the deepening operation in 1996 are presented here. the re-entry well found that the seismic event at 2340 m, prognosed from the pre-drill vertical seismic profile (vsp) as a decrease in impedance, was not base basalt and the deepened well remained within the lower series basalts. nonetheless, compressional and shear sonic logs and a density log were recorded over the full open hole interval. these allowed a firm tie to be made with the reflected wavefield from a new vsp. the sonic logs show a compressional to shear wavespeed ratio of 1.84 which is almost constant with depth. sonic compressional wavespeeds are 3% higher than seismic velocities, suggesting dispersion in the basalt flows. azimuthal anisotropy was weakly indicated by the shear sonic log but its orientation is consistent with the directions of mapped master joints in the vicinity of the well. the vsp downgoing compressional wavelet shows good persistence, retaining a dominant period of 28 ms at 3510 m depth. average vertical velocity is 5248 m/s, higher than previously reported. attenuation can largely be modelled by geometrical spreading and scattering loss, consistent with other studies. within the piled flows, the effective q from scattering is about 35. elastic layered medium modelling shows some hope that a mode-converted shear wave may be observed at moderate offsets. like its predecessor, the 1996 vsp indicates a decrease in impedance below the final depth of the well. however, it is unlikely to be basement or sediment and is probably an event within the volcanic sequence. keywords: faroe islands, lopra-1/1a borehole, basalt, vertical seismic profile, seismic attenuation _______________________________________________________________________________________________________ p.c., schlumberger cambridge research, high cross, madingley road, cambridge cb3 0el, uk. formerly: on secondment to bp, farburn industrial estate, dyce, aberdeen ab21 7pb, uk. e-mail: pafc1@slb.com i.g., fugro-jason uk ltd., unit b kettock lodge, campus 2, aberdeen science & technology park, balgownie road, bridge of don, aberdeen ab22 8gu, uk. (formerly: schlumberger geoquest, c/o bp, farburn industrial estate, dyce, aberdeen ab21 7pb, uk.) d.c., bp egypt, 14 road 252, digla, ma’adi, cairo, egypt. (formerly: bp, farburn industrial estate, dyce, aberdeen ab21 7pb, uk.) © geus, 2006. geological survey of denmark and greenland bulletin 9, 23–40. available at: www.geus.dk/publications/bull geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1923 24 the north atlantic igneous province, of which the faroe islands are a part, has been estimated to comprise 10 million km3 of intruded and extruded basaltic igneous rocks (white & mckenzie 1989). they were emplaced by the processes of rifting and sea-floor spreading which resulted in the opening of this northern part of the atlantic ocean. the basalts, which were extruded in a relatively short period of time in the early palaeogene, cover pre-existing sedimentary rocks which may well be prospective hydrocarbon traps. however, the difficulty of using reflection seismic imaging to probe beneath basalts has been recognised for some time and motivates studies into the characteristics of basalt flows which are relevant for seismic wave propagation. such studies rely on boreholes which have penetrated significant amounts of basaltic material and in which good quality geophysical logs have been recorded. the lopra-1/1a research well, on the faroe islands, is one such borehole which not only has good quality logs but also has a vertical seismic profile (vsp). the lopra-1/1a well-site is located near to the coastline on an isthmus on the southern faroese island of suðuroy. the exposed basalt sequence on the faroe islands has been divided into a lower, a middle and an upper series, each about 1 km in thickness (rasmussen & noe-nygaard 1970, 1990). the 3 km of exposed lava flows in the faroe islands are tholeiitic flood basalts whose extrusion was contemporaneous with the opening of the norwegian–greenland sea in the palaeogene. the exposed and drilled lower series lava sequence is dated to about 56–59 ma (larsen et al. 1999) and is characterised by flows with an average thickness of 20 m, thought to have originated from fissure eruption sites with a nw–se trend (rasmussen & noenygaard 1970; kiørboe & petersen 1995). the basalt flows dip to the north-east and the area around the lopra-1/1a well-site has mapped sets of near-vertical master joints trending nw–se and ne–sw. the lopra-1/1a location was selected for probing the deep basalt layers and their substrata because the ground surface is about 750 m below the top of the lower series. the lopra-1 research borehole was originally drilled to a total depth (td) of 2178 m in 1981. it was logged and suspended with most of the drilled interval left uncased. the well was surveyed by the geological survey of denmark with a zero-offset vsp and a walkaway vsp, acquired by prakla-seismos, in 1988. a refraction profile was acquired by the faroe islands natural history museum and the university of bergen in 1989. the results and interpretation are summarised by kiørboe & petersen (1995). in 1996, the well was re-entered and deepened in a number of stages by the lopra deepening consortium. the original well was extended to 3158 m kb (measured depth relative to the kelly bushing) using a larger rig with kb 16.2 m above mean sea level, retaining the original name lopra1. for technical reasons, a side-track, lopra-1a, was drilled from 3091 m kb to td at 3565 m kb. the well was logged several times during drilling, and logs were run in both the side-track and the original well. in this paper, we deal with data from a composite of the log runs in both the 1 and 1a wells, and therefore we use the name lopra-1/1a re-entry well or the lopra-1/1a deepening for the combined extensions drilled in 1996. one goal of the lopra project was the seismic characterisation of piled basalt flows. in the event that significant siliciclastic sediments were encountered beneath the basalt sequence, the consortium partners had agreed on a programme of multiple-azimuth, walkaway vsps to characterise the properties of compressional and shear wave transmission and reflection at a basalt-sediment contact as functions of both the vertical polar angle and the horizontal azimuthal angle. layered systems of high velocity contrasts, such as basalt flows, are expected to exhibit transverse isotropy with an axis of symmetry perpendicular to the layering, at wavelengths long compared to the layer thickness. such anisotropy has its fastest velocity in the direction parallel to the layering. kiørboe & petersen (1995) had reported velocities higher in the vertical direction than the horizontal and offered an explanation in vertical fractures around basalt columns, possibly in combination with the nearly vertical master joints. such vertical fractures, if aligned, would be expected to result in an azimuthal variation of seismic velocity which should be fastest in the direction parallel to the fractures. in fact, neither sediment nor basement was encountered in the well and so the borehole seismic programme was confined to a short-offset vsp and check shot survey designed to measure the short offset reflectivity at the well and to identify depths of intermediate reflectors penetrated by the bit. in particular, kiørboe & petersen (1995) had reported a reflection on the vsp at an interpreted depth of 2340 m which was thought to result from a decrease in impedance. such a reversal in impedance might have corresponded to the base of basalt/top of sediment but turned out not to be the case. further objectives of the borehole seismic analysis were to calibrate the sonic log, thereby providing a detailed velocity-depth model, and to estimate the seismic attenuation of the basalt sequence. the vsp was complemented by the acquisition over the full interval of the well of compressional (p) and shear (s) wave sonic logs, acquired in four component mode to estimate azimuthal anisotropy parameters (esmersoy et al. 1994), and a density log. the log data enabled a good well-tie to be made and allowed a modelling study to supgeus bulletin no 9 7 juli.pmd 07-07-2006, 14:1924 25 port estimates of anisotropy parameters and attenuation. this paper presents a summary of the data acquisition, compares the results with those reported in previous studies and discusses their interpretation in the context of the seismic characteristics of piled basalt flows. it draws upon previously unpublished data and reports analysed and compiled by the authors for bp, their atlantic margin partner shell, and their associates in the lopra deepening consortium. data acquisition the earlier vsp surveys (kiørboe & petersen 1995) had found evidence for a strong seismic reflector at 2340 m, just below the 1981 td, interpreted as a reversal in impedance. this event, and other deeper events seen on the vsps, were targets for the deepening of the well in 1996. further motivation was given by the discovery of methane and nitrogen at a pressure of 20 bar when the well was re-opened in 1983 (kiørboe & petersen 1995). as mentioned above, the contingency plan for the 1996 deepening project called for offset vsps at multiple azimuths in the event of finding 200 m or more of siliciclastic sediments. exploring a range of offsets was intended to evaluate the angular dependence of pto s-mode conversions in transmission within the basalt sequence, and their possible conversion back to p at the top of any basalt-sediment contact encountered. this would test the applicability to the faroese basalts of ps-mode converted imaging, as later reported by emsley et al. (1998). kiørboe & petersen (1995) had reported that the vertical p-velocity was about 10% faster than the horizontal p-velocity in the upper 800 m of the basalt sequence and appealed to fractures to explain this difference. it was hoped that the new vsp would be able to explore the nature of the vertical and azimuthal anisotropy throughout the interval of the deepened well. the location of the lopra-1/1a well on an isthmus would have facilitated the use of a marine mobile source but because of the establishment of a fish hatchery in the fjord, the 1996 vsp could not make use of a marine airgun and so twin vibroseis units were shipped to the faroe islands for land walkaway vsps in different azimuths. in the event, lopra-1a td remained in the lower basalt series and so a short-offset vsp was acquired by schlumberger on 29 october 1996 from 3510 m kb to 1320 m kb with additional checkshots up to 200 m, using the two vibroseis sources in tandem. the vibroseis units swept from 10 hz to 130 hz over 16 seconds and 2–4 sweeps were recorded at each level using a sample interval of 2 ms. in addition, checkshots were recorded using an airgun source in a water filled pit to calibrate the vibroseis transit times. the downhole tool used gimballed triaxial geophones mounted in a sensor package decoupled from the body of the tool. the deepened well was surveyed at 20 m intervals for vsp waveform processing. the cased hole section was surveyed at similar intervals to a point above the cement top where it became clear that the casing was unsupported and no longer well coupled to the formation. some logs recorded after drilling the original well were uncertain in their calibration. the sonic log was a cement bond tool with a single source-receiver pair and lacked shear sonic information (nielsen et al. 1984). since a goal for lopra-1/1a was to characterise seismic propagation characteristics for basalt flows, a set of new logs were acquired over the original open hole section, prior to setting 7-inch casing and drilling on with a 6.5-inch bit. all the drilled intervals were logged with density, pand s-sonic from a dipole shear tool in four component mode allowing estimates of azimuthal anisotropy, and a formation microscanner. the logs are not subject to petrophysical interpretation in this paper, but were used for geophysical analysis. results vsp the raw vsp traces were correlated, edited and vertically summed to produce a stacked trace at each level. the stacked vertical geophone data are displayed in fig. 1a in one-way time, static corrected to mean sea level. from the top of the vsp down to approximately 1840 m kb, the waveforms following the first arriving compressional wave are affected by borehole reverberations caused by unsupported casing in a hard rock environment with a non-attenuative fluid in the hole. the amplitude data in this interval are treated with caution, though the arrival times appear to be representative of basalt velocity. below the cement top, the stacks show good waveform consistency from level to level. clearly visible are: a. down-going multiples (parallel to, but later than, the first arrivals); b. up-going primary reflections, both within and below the drilled interval. these are characterised by an almost linear moveout of equal slope but opposite sign to that of the first arrivals; c. suggestions of down-going shear energy, with a linear moveout greater than that of the first arrivals; d. a weak tube-wave (visible only above the cement top). geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1925 26 the vertical one-way time from mean sea level to the deepest vsp level at 3510 m kb (equivalent to 3478 m true vertical depth sub-sea: tvdss) is 663 ms, corresponding to a compressional velocity of 5248 m/s over the interval. this is higher than values for basalt velocity reported elsewhere in the literature (e.g. planke & cambray 1998) and higher than the 4.35 km/s average velocity estimated from the first logs run in lopra-1 (nielsen et al. 1984). this is partly due to the high velocity dolerites and partly due to the relatively thick flows in lopra1/1a, giving a higher thickness ratio of fast, flow-centre material to slower, flow-boundary material. as discussed later, the average compressional wavespeed from the sonic log is slightly higher, supporting the vsp observation. the trace scaling in fig. 1a is constant for all levels, revealing the total amplitude loss with depth in the first arrivals. although the vsp interval, from 3510 m kb to 1320 m kb, comprises 2174.5 m tvd of stacked basalt flows, the amplitude loss from geometrical spreading, scattering and attenuation still leaves a good level of signal above the noise floor in the deepest section. fig. 1b shows the y-component stacks (y is the horizontal component tangential to the borehole wall) on which the shear energy is most evident. the strongest arrivals are down-going direct shear arrivals, generated by the vertically polarised vibrator acting on the rigid surface. the average shear velocity across the logged interval is about 2900 m/s, resulting in an average vp/vs ratio of 1.8, which is in good agreement with the value of 1.84 + 0.01 (one standard deviation) estimated below from the sonic log regression of vs upon vp. also visible on the y-component stacks are down-going mode-converted shear events, e, generated by impedance contrasts crossed by the drill-bit. these events are parallel to and earlier than the direct shear arrival, c, but originate at the time and depth of the direct compressional wave’s encounter with the mode-converting impedance contrast. these arrivals have a frequency content similar to the associated p-wave and higher than the frequency of the direct shear wave due to the lower cumulative attenuation. finally, some weak up-going reflected shear events can be seen. no further processing of the shear arrivals has been performed. the vsp p-wave data were processed through a workflow comprising trace editing, vertical stacking at each depth level, correction to mean sea level, spherical divergence correction, up/down wavefield separation and deconvolution to zero-phase wavelets with bandwidths of 10–70 hz and 10–40 hz. the former maximises the resolution at the cost of more high frequency noise, while the latter minimises residual noise at the expense of bandwidth. correlations of the 10–70 hz up-going wavefield are presented and discussed in the well-tie section. geophysical log data in this section we examine the 1996 log data and their correlations to infer the seismic properties of the basalt and to compare with the properties of similar basalts penetrated in hole 917a of the ocean drilling programme, as reported by planke & cambray (1998). the pand s-sonic and density logs, spliced from the fig. 1. a: lopra-1/1a vsp: stacked vertical component geophone data. b: lopra-1/1a vsp: stacked tangential component geophone data. labelled phases are described in the text. 0.2 0.4 0.6 0.8 1.0 1.2 1.4 1.6 3510 1340 measured depth kb (m) vertical component 0.2 0.4 0.6 0.8 1.0 1.2 1.4 1.6 3510 1340 measured depth kb (m) r ec or de d tim e (s ) tangential component b b a c c ed a b geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1926 27 kb, the core measurements are either the same as the log values or a little faster, consistent with dispersion of the higher frequency core measurement compared to the sonic log. discussion layer-induced anisotropy backus upscaling applied to a stack of isotropic layers predicts a transversely isotropic model having a symmetry axis perpendicular to the layers. the upscaled elastic parameters can be recast in terms of thomsen’s (1986) weak anisotropic parameters. figure 3 shows the modelled values of epsilon (ε), and delta (δ), after upscaling the logs to 3 m. the parameters are given by the relations: and where and ρ, vp and vs are the backus-upscaled log values of density, compressional wavespeed and shear wavespeed. epsilon is the measure of axial compressional anisotropy, while delta controls the off-axis behaviour of the phase velocity near to the vertical. both enter the following relation from thomsen (1986) governing the compressional phase velocity behaviour as a function of angle θ to the (vertical) axis of symmetry: (c13 + c33)(c13 – c33 + 2c55) 2c33 (c33 – c55) several logging runs made in the 1 and 1a wells, are shown in fig. 2 on a measured depth scale relative to kb. the well is close to vertical, so no true vertical depth corrections have been made for this part of the analysis. the logged section comprises 3371 m, with the dipole shear tool logged in four-component mode together with the density tool. the logs are of good quality and allow the identification of many subaerially emplaced flows over much of the section. figure 2 shows the pand s-velocity and density logs after re-sampling the data to 3 m using a backus average (backus 1962; folstad & schoenberg 1992). there are wide variations in pand s-velocities, in an asymmetric, quasi-periodic manner, though the mean velocities are very consistent over the well. the boundaries of each flow are characterised by a shift to lower velocities and density, caused by the formation of vesicles at the top and base of the flow (planke & cambray 1998), and possibly in places by weathering, alteration and rubble. the main exceptions to this character are the high velocity dolerite intrusions, encountered at about 600 m kb and 770 m kb, and the zone of almost constant velocity and density between about 2600 m kb and 2900 m kb, corresponding to a thick, hyaloclastite sequence. the dots near the compressional velocity log (vp in fig. 2) mark depths and values of ultrasonic measurements of p-wave velocity made by geus on a number of core samples. the measurements were taken on 25 mm core samples, pressure-saturated with distilled water and using piezo-electric transducers of centre frequency 1 mhz (2380 m core) and 2.5 mhz (all other cores) at room temperature and pressure. although the scale of the display makes a visual match difficult, it can be seen that there is generally a good correlation between the core measurements of p-wavespeed and the log data. with only two or three exceptions, at 2218 m kb, 2455 m kb and 2972 m fig. 2. p-velocity, s-velocity and density logs, upscaled to 3 m sample interval using a backus average (backus 1962) and displayed in measured depth relative to kb. core points are plotted from data supplied by geus for comparison. 7000 6000 5000 4000 3000 2000 1000 0 ve lo ci ty ( m /s ) d en si ty ( g/ cm 3 ) measured depth kb (m) 0 350030002500200015001000500 3.2 3.0 2.8 2.6 2.4 2.2 core data points geus core lab vp vs ρ ε = c11 – c33 2c33 δ = c33 = pvp , c55 = pvs , c13 = c33 – 2c55 2 2 geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1927 28 vp (θ) ≈ vp (0)(1 + δ sin2 (θ) cos2 (θ) + ε sin4(θ)) in fig. 3, the anisotropic parameters are greatest where the logs show the most variance over the 3 m averaging interval. they indicate only the layering component of the anisotropy since we have no data on the intrinsic anisotropy of the basalt. however, delta is small with a mean of zero while epsilon is positive with most values less than 0.05 and almost all values less than 0.1. we found a similar range of epsilon and delta when the averaging interval was increased to 10 m. this means that the modelled effect of layering is to produce horizontal compressional wavespeeds around 5% faster than vertical wavespeeds, a prediction which contradicts the observation by kiørboe & petersen (1995) of vertical compressional wavespeeds being about 10% higher than horizontal compressional wavespeeds in the upper 800 m of the basalt beds. kiørboe & petersen (1995) appeal to vertical fractures associated with columnar basalts and master joints to explain their observation, a point which we discuss below. azimuthal anisotropy direction aligned vertical fractures would be expected to result in azimuthal anisotropy. to test for this in lopra-1/1a, the dipole shear sonic tool was logged in four-component mode whereby both of the two orthogonal dipole receiver arrays recorded signals from each of the two orthogonal dipole sources, energised sequentially. the data were processed for the presence of fast and slow shear waves corresponding to shear propagation along the borehole with polarisations parallel and perpendicular to the assumed fractures. one output of the azimuthal processing is a log of the azimuth of the fast shear wave. the direction of the fast shear wave is found by determining the shear waveform rotation which minimises the cross-energy, for example the energy recorded by the y-polarised receiver from the x-polarised source (esmersoy et al. 1994), then picking the faster of the two shear estimates. the difference in cross-energy between its maximum and minimum excursions as a function of rotation angle is a measure of reliability of the anisotropy estimate. in figs 4a and 4b we plot the fast shear direction results in different depth intervals, where the vector from the centre of the plot to a given depth point has an azimuth corresponding to that of the fast shear polarisation and a magnitude corresponding to the cross-energy difference normalised by the total energy. in order to display only the most reliable estimates of the fast shear direction, the estimates were windowed according to cut-off values of calliper reading, the estimated error in azimuth, the anisotropy estimate and the normalised cross-energy difference (cut-off values indicated on the figure). the plots are radially symmetric through the origin because of the 180° ambiguity in determining azimuth. in fig. 4a, from 511–540 m kb, we see that there is a well-defined fast shear direction at n31°e, broadly consistent with the ne–sw strike of one of the two mapped master joint sets (kiørboe & petersen 1995, fig. 1). below 540 m kb, the fast shear direction rotates through 100° to n131°e, although this direction is less well defined. in fig. 4b, there is a fairly consistent fast shear direction of about n144°e in the interval 2991–3502 m kb, but with a larger scatter in azimuth. the cross-energy cut-off value in fig. 4b has been reduced, compared to fig. 4a, to capture more estimates. the estimates of the fast shear direction deeper than 540 m kb are consistent with the nw–se strike of the second mapped joint set in kiørboe & petersen’s (1995) fig. 1. from the earlier vsp and refraction profile data, kiørfig. 3. log of thomsen’s (1986) weak anisotropic parameters epsilon and delta, modelled from the backus upscaling of the 15 cm logs to 3 m, assuming isotropic individual layers. the compressional sonic velocity is shown for correlation, also upscaled to 3 m.d el ta a nd e ps ilo n measured depth kb (m) ve lo ci ty ( m /s ) 7000 6000 5000 4000 3000 2000 1000 0 0 350030002500200015001000500 0.10 0.05 0.00 –0.05 vp epsilon delta geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1928 29 boe & petersen (1995) reported vertical vp values about 10% higher than horizontal vp in the upper 800 m of the section around the well and suggested that a combination of vertical, columnar fractures and nw–se master joints crossing the ray-paths in the refraction profile may be responsible. the present analysis shows little azimuthal anisotropy on the dipole shear log and, while the fast shear direction appears to be n131°e below 540 m kb, there is a well-established, fast shear direction of n31°e from 511–540 m kb. if the shear azimuthal anisotropy results from fractures or stress, then the compressional wavespeed anisotropy should follow the same directions. formation microscanner data from the lopra deepening project show the presence of sub-horizontal conductive features, suggestive of horizontal fractures, although there is a question as to whether these are natural or drilling induced. the only conventional core (2380 m kb) shows cemented, horizontal fractures (l. kiørboe, personal communication 1996) but few vertical fractures were observed, consistent with the weak azimuthal anisotropy observed in fig. 4. a: fast shear direction estimated from the dipole shear log over the depth interval 511–1010 m kb. the log was recorded in four component mode and rotated to minimise the cross-energy. two orthogonal directions are evident. the better-defined direction is over a fairly short interval from 511–540 m kb, consistent with one of the mapped master joints, and parallel to the offset seismic surveys described in kiørboe & petersen (1995). b: fast shear direction estimated over the interval 2991– 3530 m kb. the fast direction is less well defined than in the shallower interval but is consistent with mapped joints. user depth 511–1010 m data extent 174.346–2121.4 m cutoff values max. calliper (inch): 11 min. itt-based anisotropy: 1 max. azimuth error (deg): 10 min. cross-energy difference: 20 450 540 630 720 810 900 990 1080 major tectonic axes so ut h– n or th ( % ) 80 40 0 –40 –80 80400–40 west–east (%) –80 d ep th ( m ) a 3600 3500 3400 3300 3200 3100 3000 2900 20 10 0 -10 -20 b b user depth 2991–3530.19 m data extent 2991–3502.91 m cutoff values max. calliper (inch): 11 min. itt-based anisotropy: 1 max. azimuth error (deg): 10 min. cross-energy difference: 1 major tectonic axes so ut h– n or th ( % ) 20 10 0 –10 west–east (%) 30 –20 –30 d ep th ( m ) 20 10 0 –10 –20 –30 b geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1929 30 the dipole shear log. in addition, the layer-induced anisotropy modelled by upscaling the log data suggests that horizontal compressional wavespeeds should exceed vertical wavespeeds by around 5%. can we resolve the apparent contradiction? possible explanations include: (1) the high-wavespeed, dolerite intrusions seen on the logs down to 850 m kb affect the vertical velocity locally; (2) the layeringand fracture-induced anisotropy revealed by the well logs is not representative of the volume sampled by the vsp and earlier refraction profile; (3) the unreversed refraction profile sees low apparent velocities because of the basalt flows that dip 3–7° to the north-east; and (4) there is a mis-correlation of events between the refraction profile and the earlier vsp. kiørboe & petersen (1995) reported that a 700 m refraction profile shot towards the wnw from the 1988 vsp source point through the well yielded an apparent vp of 4.8 km/s, the same as that observed over the offset range 2.3–14 km on the long refraction profile. the azimuth differences of these two profiles weakens the case that structural dip or aligned master joints affect the velocities. that the fractures in the core sample are cemented also discounts the possibility that open, vertical, hexagonal fractures are prevalent, although we cannot disprove this possibility because of the small volumes sampled by core and borehole logs. our preferred explanation is that the geometry of the dolerite intrusions increases locally the estimate of the vertical velocities at the well. further support for this hypothesis comes from the well-tie, described below, where there is evidence that the vsp sees a higher velocity than the sonic tool in the shallow section, possibly due to the dolerite intrusions providing a vsp ray-path away from the borehole that is faster than that seen by the sonic tool along the borehole. kiørboe & petersen (1995) also found that the velocities derived from the 1988 vsp were too fast to simulate the arrival times on the longer offset data. log-derived seismic characteristics and well-tie in this section we develop further the seismic properties of the stacked basalt flows in lopra-1/1a as modelled from the log data. figure 5 displays the vp/vs ratio and the normal incidence, two-way geometrical spreading, computed using newman’s (1973) relation: for the 3 m backus-averaged data since there is no expected dependency on sample interval. in this relation, dtwoway is the two-way loss, relative to the amplitude at 1 m from the source, of a seismic wave propagating down through a 1d stack of i layers, each of thickness di and interval velocity vi, and back to the surface again. v1 is the velocity of the first layer and in this case is taken to be 2792 m/s. because the thickness of the superficial layer between the source and the top of the sonic log is 177.67 m, the two-way loss at the top of the sonic log is: dtwo–way = – 20log(2 × 177.67) = – 51.01 db re 1 m figure 5 shows the modelled geometrical spreading in the basalt flows, compared to the spherical loss in a uniform medium for reference. the extra two-way geometrical spreading due to the velocity variations is around 5.5 db for a reflector at the total depth of the well, resulting in a spreading loss of 82.5 db re 1 m. the vp/vs ratio from the 3 m averaged logs (fig. 5) is almost constant over the well at 1.85, with a standard deviation of 0.04. we also cross-plotted the log values of vp and vs, upscaled to 1 m to reduce fluctuations, using a backus average. figure 6 shows the cross-plot of vp to vs with a histogram of each variable indicating the number of data points falling into velocity bins of 100 m/s for vp dtwo–way = 2 σdi vi vl fig. 5. plot of logged vp/vs ratio and short-offset geometrical spreading computed using newman’s (1973) relation for both the logged velocities and a constant velocity medium. 3 m averaged logs. –85 –90 –75 –65 –55 v p/ v s ra tio measured depth kb (m) lo ss es ( db r e 1m ) 0 3600300020001000 2.8 2.6 2.4 2.2 2.0 1.8 1.6 1.4 –80 –70 –60 –50 geometrical spreading spherical spreading vp/vs geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1930 31 and 55 m/s for vs. the bin-widths are scaled approximately by the vp/vs ratio, for which a value of 1.84 was obtained by linear regression with a standard deviation of 0.01. the similarity of the histograms supports the linearity of vp and vs over such a large range of values. this linearity is consistent with the observation of planke & cambray (1998) who studied subaerially emplaced lava piles sampled in hole 917a of the ocean drilling programme. the absence of a trend, with the possible exception of the hyaloclastite interval from 2600–2850 m kb, suggests that there is no significant compaction effect with depth. we computed two-way, plane-wave, transmission losses for the 15.24 cm logs as well as for logs backus-averaged to 30 cm, 50 cm, 1 m, 3 m, 5 m and 10 m (actually, the nearest multiple of 15.24 cm to these values). the transmission losses are the 1d, ray-theoretical loss from a unit amplitude plane-wave source pulse which is partially transmitted and reflected from each interface in its twoway path from the surface to a given reflector and back to the surface again. it is computed from the well-known relation tl = π (1 – r2 i ) where ri is the plane-wave, normal-incidence, reflection coefficient at the i-th boundary ri = and where ρi and vpi are the density and compressional wavespeed of the i-th layer. in this simple model, there is no intrinsic loss. instead, the loss is due to progressive scattering and consequent removal of energy from the firstarriving, primary pulse. the results are displayed in fig. 7 along with the vp log for correlation. losses over the full interval vary non-linearly from 15 db (10 m log) to 73 db (50 cm log), because of the non-linear dependence on the reflection coefficient. the highest loss is seen at the 50 cm sampling. as suggested by o’doherty & anstey (1971), an impedance gradient represented by a single large reflection coefficient has more effect on the transmission losses than several smaller coefficients, so very fine sampling (15 cm) can produce a smaller loss than the upscaled logs, as seen here. however, blocking intervals larger than bed thicknesses progressively fail to represent beds at all and so will predict less loss. while there is the expected strong dependency of transmission losses with sample interval, all the curves show broadly similar features, with the lower contrast, hyaloclastite interval below 2600 m kb exhibiting reduced loss compared to the section of high contrast flow units, which resembles the cyclic example in o’doherty & anstey (1971) and which is expected to have a high scattering loss. the combined effects of modelled normal-incidence spreading and plane wave transmission losses range from 97.5 db to 155.5 db re 1 m, where the wide margin results from the uncertain effects of the transmission losses. we shall return to this point in the analysis of the vsp amplitudes. however, we emphasise that the loss estimates here are those due to near-normal propagation in a 1d medium. as offset increases, geometrical losses increase rapidly due to refraction effects at high velocity contrast boundaries. similarly, transmission coefficients vary strongly with incidence angle at high contrast boundaries. the cone of forward propagation of the compressional wave is limited by critical angle effects and offers opportunities for mode conversions to be observed over a wider range of incidence angles and offsets. the vsp first arrival travel times were picked to derive fig. 6. vp–vs cross-plot of 1 m backusaveraged logs with histograms and linear regression. the histograms show the frequency distributions of vp and vs data gathered in bins of 100 m/s and 55 m/s respectively, to maintain the same number of bins for each variable. the regression yields a vp/vs ratio of 1.84 with a standard deviation of 0.01. the similarity of the histograms supports the choice of a linear regression. ρi + l vpi + l – ρivpi ρi + l vpi + l + ρivpi vp (m/s) v s (m /s ) frequency fr eq ue nc y vp histogram vs histogram vs = 0.5437 vp – 16.647 r2 = 0.9512 0 100 200 200 100 0 3500 2500 1500 3500 4500 5500 6500 7500 geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1931 32 a seismic two-way time to true vertical depth curve. the time–depth curve from the vsp was used to ‘calibrate’ the tvd-corrected sonic log, provide a fine-scale, seismic time–depth interpolation and to generate a synthetic seismogram. a drift curve was computed by integrating the compressional slowness log between the vsp time-depth pairs and taking the difference between the vsp and sonic interval times. we observed positive drift (vsp interval time greater than sonic) over most of the logged interval. this implies a lack of environmental effects on the sonic (which usually increase the sonic slowness) leaving the normal dispersion effect between seismic and sonic frequencies as the dominant mechanism. negative drift (sonic interval time greater than vsp) was seen in the shallow part of the well, above 800 m kb. this would be normal in a sedimentary sequence, indicating a washed-out hole or altered, unconsolidated formations. however, in the basalt environment, formation alteration is unlikely; even though the well shallower than 800 m kb is often out of gauge, it is not different from the rest of the borehole drilled in 1981 which displays positive drift. possible explanations of the negative drift are unreliable vsp picks due to shallow, unsupported casing interference (the compressional wavespeed of the formation is similar to that of the casing extensional mode), or possible refraction along the high velocity (6.5 km/s) dolerite intrusions. if the vsp wave front refracts along the intrusions such that the borehole does not form the vsp raypath, then the vsp interval time will be less than the sonic interval time. in any case, the picks above 800 m have been interpreted in such a way as to avoid undue, and possibly unrealistic, correction to the sonic. the corrected compressional sonic log was used to convert the shear sonic and the density logs from a common depth scale to a common two-way time scale which is assured to match that of the vsp. note that this implicitly changes the vp/vs ratio, since the vsp was not used to correct the shear sonic slowness values. deeper than 800 m, the compressional slowness drift is about 15 ms in 2200 m, or 6 µs/m. however, the average interval slowness is about 190 µs/m, so the drift correction is reasonably significant at 3.4% in slowness. care was taken in choosing the correction points to avoid introducing false reflection events. the time-based logs drove a 1d model of equal timethickness layers. the resulting primaries-only, reflection coefficient sequence without transmission losses was convolved with a 40 hz, zero-phase, ricker wavelet to create a synthetic seismogram. the synthetic was spliced into the vsp up-going wavefield along the two-way time–depth curve to facilitate event correlation (the right-hand panel of fig. 8). the vsp has been waveshaped to a zero phase wavelet of bandwidth 10–70 hz and both the synthetic and vsp traces are displayed in reverse seg polarity, so that an increase in acoustic impedance with depth is displayed as a black peak. figure 8 shows the correlation from a two-way time (twt) of 650 ms, where several events can be seen to tie in time and in character, resulting in an unambiguous correlation. in an igneous province, a tie of this quality is relatively unusual and suggests that the lateral variability of the basalt flows is mild, at least over the extent of a vsp fresnel zone which is several tens to hundreds of metres, depending on the elevation of the vsp tool above the reflector. residual ringing from the uncemented casing is visible in the up-going vsp wavefield on the right edge of the section. correlations from the well-tie using fig. 8 we discuss the correlation to the lithological summary taken from the lopra-1/1a end of well report (ewr), subsequently modified by r. waagstein (personal communication 2001). the displayed interval from 650–1430 ms twt shows vp, vs, density and poisson’s fig. 7. modelled plane-wave, normalincidence, two-way transmission losses as a function of log sampling, together with vp log for correlation. losses are nonlinear with sampling but show generally similar behaviour with high contrast layering resulting in the greatest scattering loss. 10 m 5 m 3 m 15 cm 1 m 30 cm 50 cm lo ss es ( db ) –80 –60 –40 0 –20 measured depth kb (m) 0 3600300020001000 8000 6000 4000 v p (m /s ) vp geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1932 33 ratio logs. at the left edge of the display is a non-linear scale in true vertical depth extending from 1714 m to td at 3533 m tvdss, which corresponds to 3565 m kb. horizontal lines mark particular correlations between the well logs, the lithology summary, the synthetic seismogram and the vsp. these correlations are summarised in table 1 and further comments relating to the pre-drill targets are made below. from 882–912 ms, the vsp shows a package of strong reflections, with a white trough at 894 ms, seen on the pre-drill vsp at 920 ms and identified as one of the target horizons for the well. this event was prognosed by kiørboe & petersen (1995) at 2.34 km, 162 m below the original td of lopra-1 and interpreted as a decrease in impedance. in fact, the reflection package comprises at least two near-tuning events. a sharp increase in impedance at 882 ms (2361 m kb, 2345 m tvdss) corresponds to the top of a massive basalt flow that gives rise to a black peak on the vsp. the trough associated with the sidelobe to this wavelet reinforces a broad, weak trough caused by the decrease in impedance at the base of the flow around 894 ms (2401 m kb, 2385 m tvdss), which reflects the highly amygdaloidal top of an underlying flow. the reflection package is seen more strongly on the vsp than on the synthetic seismogram, possibly due to lateral heterogeneity near the well. another package of strong events was drilled above td from 1274–1340 ms. during drilling, it was hoped that these reflections, seen at 1350 ms on the pre-deepening vsp (kiørboe & petersen 1995), might mark either siliciclastic sediment or basement. however, the sharp drop in impedance at 3427 m kb (3397 m tvdss) at the base of a thick, massive basalt bed results in the trough at 1294 ms seen on both the vsp and the synthetic, which correlates with another hyaloclastite sequence with a thin tuff at the top (r. waagstein, personal communication 2001). another drop in impedance was drilled at 3512 m kb (3480 m tvdss), which also corresponds to a strong white trough at 1328 ms on both the synthetic and the vsp. it probably corresponds to the event at 1350 ms seen on the pre-drill vsp. it displays moveout to earlier times with decreasing geophone depth, indicating dip of the beds. just below td, at 1412 ms on the vsp, is a persistent, large-amplitude, symmetric, white trough, which may be fig. 8. correlation of the vsp up-going wavefield with the log-derived, reflection coefficient sequence convolved with a 40 hz zero-phase ricker wavelet. the vsp wavefield has been waveshaped to zero phase over the bandwidth 10–70 hz. the vsp and synthetic are correlated with the time-based logs and lithostratigraphy (r. waagstein, personal communication 2001) as listed in table 1. lithology summary basalt flows hyaloclastites massive flow hyaloclastites hyaloclastites + basalt beds hyaloclastites without basalt beds hyaloclastites with basalt beds massive flow hyaloclastites hyaloclastites with minor basalt beds 1743 1797 1848 1900 1957 2007 2065 2122 2177 2234 2289 2342 1402 2454 2502 2550 2600 2649 2698 2748 2796 2844 2894 2995 2945 3048 3098 3151 3204 3256 3307 3360 3461 3413 3514 tvd ss d ep th (m ) 3000 5000 7000 p-wave (m/s) 2000 3000 s-wave (m/s) 1.5 2.0 2.5 3.0 0 density (g/cm3) 0.30 0.10 0.50 poisson t im e (m s) sp of f 700 800 900 1000 1100 1200 1300 700 800 900 1000 1100 1200 1300 1400 synthetic + vsp 3 7 11 16 21 26 31 36 41 46 51 56 61 66 71 76 81 86 91 96 102 105 total depth = 3565m md kb = 3533 tvd ss geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1933 34 t ab le 1 . v sp e ve nt s v sp e ve nt t op o f v sp t op o f pa ck ag e of s tr on g re fle ct io ns ; sh ar p in cr ea se in im pe da nc e d ec re as e in im pe da nc e ba se o f p ac ka ge o f s tr on g re fle ct io ns m od er at e re fle ct io ns t op o f l ow r ef le ct io n in te rv al ba se o f l ow r ef le ct io n in te rv al t op o f p ac ka ge o f s tr on g re fle ct io ns in cr ea se in im pe da nc e bo tt om o f v sp d ec re as e in im pe de nc e, d ip pi ng b ed bo tt om o f p ac ka ge o f s tr on g re fle ct io ns t d ( to ta l d ep th o f l op ra -1 /1 a ) si gn ifi ca nt d ec re as e in im pe da nc e * f m s lo g in te rp re ta tio n (r . w aa gs te in , p er so na l c om m un ic at io n 20 01 ) † k iø rb oe & p et er se n (1 99 5) ba sa lt flo w s h ya lo cl as tit es t op m as si ve b as al t flo w (2 36 1– 24 01 m ) a m yg da lo id al b as al t flo w (2 40 1– 24 17 m ) h ya lo cl as tit es w ith b as al t be ds h ya lo cl as tit es w ith n o ba sa lt be ds h ya lo cl as tit es w ith b as al t be ds m as si ve b as al t be d (3 39 6– 34 27 m ) to p hy al oc la st ite s (0 .9 m t uf f a t to p) ba se m as si ve b as al t be d (3 50 4– 35 12 m ); hy al oc la st ite s w ith m in or b as al t be ds ( 35 12 –? m ) 49 6 86 5 88 2 89 4 91 2 91 2 98 0 10 88 12 74 12 82 12 94 13 26 13 28 13 40 13 47 14 12 13 20 23 19 23 61 24 91 24 50 26 16 28 82 33 96 34 27 35 10 35 12 37 32 t w t m ill is ec . 92 0 13 50 d ep th be lo w k b m et re s 23 98 24 43 26 12 28 80 34 26 35 12 lo w er b as al t se ri es pi llo w la va s er ie s pi llo w d eb ri s se ri es + ba sa lti c tu ff 1 + tu ffa ce ou s sa nd 1 ba sa lti c tu ff 2 + ba sa lti c sa nd 1 + t uf fs 1 –2 tu ffs 3 –9 ba sa l p ar t of t uf f 9 tu ff 10 pr eba sa lti c tu ff se ri es b (t uf fs 1 1– 12 ) d an sk o lie og g as pr od uk tio n & d an sk o pe ra tø rs el sk ab ( 19 97 ) v sp b ef or e de ep en in g† li th ol og ic al u ni t* t w t m ill is ec . d ep th be lo w k b m et re s t v d ss m et re s d ep th be lo w k b m et re s 23 40 li th os tr at ig ra ph ic u ni ts lo pr a1/ 1a v sp sh ar p de cr ea se in im pe nd en ce ba se m as si ve b as al t be d; 35 65 13 04 23 03 23 45 23 85 24 34 26 00 28 64 33 66 33 97 34 78 34 80 35 33 37 00 geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1934 35 interpreted as a significant drop in impedance. this is about 86 ms, or 222 m, below the deepest vsp level at 3510 m kb (3478 m tvdss) and therefore 167 m below the final td of 3565 m kb (3533 m tvdss), assuming an extrapolation velocity of 5150 m/s. the event is as tantalising as that seen at 920 ms on the pre-deepening vsp, which was drilled some 162 m below the original td of lopra-1. however, similar, though weaker, ‘soft kicks’ seen on the seismic data have turned out to be due to contrasts within the volcanic pile and the reflection at 1412 ms twt is unlikely to be basement, so it is probable that more volcanic sequence lies below thepresent td. in contrast with the event at 1328 ms, this strong arrival below td shows little moveout and suggests low dip. vsp loss estimates and amplitude modelling as noted above, ray-theoretical estimates of transmission losses are strongly dependent on the log sampling, so we derivedloss estimates from the vsp downwavebefore turning to full waveform modelling to simulate the observed propagation effects. the contoured power spectrum by depth is shown in fig. 9, after windowing the first 150 ms of data from the 2 ms sampled downgoing wavefield. two points are apparent: (1) there is a low-frequency roll-off to the data caused by the start frequency of the vibrator sweep at 10 hz, (2) there is a smooth decline in frequency content with depth. we estimated the root mean square (rms) amplitudes within a window of 150 ms about the vsp first breaks, corrected the amplitudes for geometrical spreading using the spreading loss curve displayed in fig. 5 and plotted the results in fig. 10 together with the ray-theoretical transmission losses at 3 m, 5 m and 10 m sampling. since the amplitude of the vsp top level is arbitrary, we matched the slopes of the transmission loss curves by eye. the vsp amplitude decay curve, after spreading correction, shows a character similar to the modelled transmission losses, although, given our earlier comments on the unreliability of transmission losses, the match to the transmission loss curve at 5 m sampling is probably coincidental.however, the change in slope around 2500 m kb on the modelled curves, which is due to the transition from high contrast basalt beds to low contrast hyaloclastites, is also evident on the vsp. by reciprocity, the transmission seismogram going back up through the basalt sequence is the same as that going down through the sequence. hence, the reflection seismogram should be the time-delayed, one-sided correlation of the down-going wave. strictly speaking, the two-way loss estimates should be made after convolving fig. 9. contoured power spectrum of the energy in a 150 ms window around the down-going first arrivals in the vsp. the upgoing wavefield was removed first. the contour levels are db down from the peak value. fig. 10. losses from rms amplitudes estimated in 150 ms window around the first arrivals in the vsp, after correcting for geometrical spreading using the logderived spreading loss estimates from fig. 5. the first vsp level amplitude is arbitrary and has been selected visually to overlie the upscaled log curve of modelled transmission losses with the most similar slope. regression lines have been fitted over the vsp intervals indicated to estimate effective q for a wavelet dominant frequency of 36 hz. 200 160 120 80 40 0 1320 1720 2120 2520 2920 3320 –30 –20 –20 –10 –40–40 –30 –20 –20 –10 fr eq ue nc y h z depth kb (m) geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1935 measured depth kb (m) lo ss es ( db ) 0 3600300020001000 –45 –25 –20 –15 –10 –5 0 regression (i) q = 35 regression (ii) q = 112 transmission losses from upscaled logs vsp estimated losses 10 m 5 m 3 m –40 –30 –35 36 the down-going wave with itself. since the amplitude effect of cascaded filters is to take the product of their gain functions, we have approximated the amplitude effect of the two-way propagation by doubling the db loss estimated from the one-way measurement of rms time-domain amplitudes. because of the change in slope of the transmission loss curves around 2500 m kb, we estimated two effective q factors to represent the scattering loss. over the interval of periodic layering, from 1350 m kb to 2450 m kb, we fitted a linear trend and estimated losses corresponding to an effective q of 35 for a dominant period wavelet of 27.5 ms. the 95% confidence limits on the regression map to a range in effective q of 32 to 40. from 2450 m kb to 3510 m kb, the loss curve is reduced and a linear regression resulted in a loss corresponding to an effective q in excess of 100. the inverse correlation of effective q values with the change in impedance contrasts further supports the inference that the volcanic sequence has low intrinsic loss and that its attenuation is principally due to scattering. this is consistent with the results of other, similar studies (e.g. pujol & smithson 1991) which found that spreading and scattering losses could account for observed vsp amplitude behaviour in basalt sequences. evidently, we can find a log sample rate such that the ray-theoretical losses match the vsp loss estimates. however, given the large variability in the transmission losses with sampling, we undertook further analysis to test the hypothesis that scattering loss is the dominant mechanism. we estimated spectral ratios of windowed vsp downwave traces at 2420 m, 2980 m, 3044 m and 3500 m kb, using the downwave at 1880 m kb as a reference. the spectral ratios are plotted in fig. 11 with the amplitude spectrum of the reference level. over the rather limited frequency interval of the strongest signal (17–40 hz), the slopes of the spectral ratios are effectively the same and rather flat. in a lossy medium, the slopes are modelled by 20 π∆z qc where c is the average velocity over the interval ∆z. the four slopes in fig. 11 correspond to depth intervals of 548 m, 1100 m, 1164 m and 1620 m, but there appears to be little variation of slope with the depth interval. the real vsp amplitude loss is not well modelled by frequencydependent attenuation. nevertheless, the variation in spectral power with depth in fig. 9 does show a loss of high frequencies with depth, mainly in the shallower part of the section, with little apparent bandwidth change over much of the deeper part of the volcanic sequence. we therefore modelled the propagation of a vsp pulse generated at ground level (gl) through a 185 m uniform layer on top of the pile of basalt flows logged from 185 m below gl and emerging into a uniform half-space at 2160 m below gl (fig. 12). the 1975 m interval was modelled using a full-elastic, 1d modelling code based upon the kennett algorithm (kennett 1974, 1983) and developed at schlumberger cambridge research. the 15 cm log data were upscaled to 3 m by backus averaging to reduce the computational cost. equivalent medium averaging, using windows of approximately 1/20th of the dominant seismic wavelength, provides a convenient method of upscaling logged data to allow efficient elastic waveform modelling, while retaining fidelity of both the travel time and amplitude information of the log scale model (folstad & schoenberg 1992, 1993). propagation was modelled both with and without multiple scattering, with 3d geometrical spreading from a point source. the injected wavelet at the top of the stack is a zero-phase, 60 hz ricker wavelet. (the wavelet appears not to be zero phase because of near field effects.) at the base of the stack, the escaping wavelet modelled without multiples is zero phase, but its amplitude has been diminished, undergoing a one-way loss of 46.9 db due to fig. 11. spectral ratios computed from four vsp levels relative to a reference level at 1880 m kb. the amplitude spectrum of the reference trace is also shown to indicate the signal frequency band. – log10e sp ec tr al a m pl itu de frequency (hz) sp ec tr al r at io ( db ) 0 10 20 30 40 50 60 400 800 1200 0–25 –20 –15 –10 –5 0 5 2420 2980 3500 3044 1880 spectrum “ “ “ “ geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1936 37 the combination of one-way geometrical spreading and transmission losses. the escaping wave modelled with all internal multiples has been phase rotated within the 40 ms analysis window and has a higher rms amplitude than the wavelet propagated without multiples, because shortperiod multiples have served to boost amplitudes within the analysis window. the rms amplitude of the escaping wave with multiples is –37.5 db relative to the input wavelet at the top of the stack. however, energy has been removed from the front of the wavelet, its trough-to-trough duration has increased and a coda has developed. these effects are due to scattering by the high-contrast layering. we ran a further model, with multiples, to study the frequency-dependent effects of multiple scattering by the periodic layering by using a wide-band source signature defined in the frequency domain from dc to the nyquist frequency at 125 hz, corresponding to 4 ms sampling (in contrast to the real vsp which was acquired at 2 ms sampling). we also altered the model by adding a lower halfspace simulating a massive sand unit and placing the deeper receiver at 300 m into the sand below the base basalt which was again modelled at 2160 m. the fourier domain amplitudes of the deepand shallow-receiver down-going compressional waves are displayed in fig. 13a, where the low-pass filtering effect of the basalt sequence can be seen, with quite deep notches appearing from 39 hz, although higher frequency peaks also appear, such as that at 85 hz. the spectral ratios of the wide-band synthetics are shown in fig. 13b with a linear regression to 80 hz, avoiding the higher frequency side-lobes. the regression provides an effective q estimate of 32, with 95% confidence limits of 26 and 42, which is consistent with the time domain estimate of 35 from the real vsp over the interval of basalt flows shown in fig. 10. however, we note that while the real vsp displayed little or no frequencydependent loss over the interval from 1880–3500 m kb, the full elastic modelling does support the assertion that the amplitude loss can be modelled by scattering and geometrical spreading alone. loss estimates from the literature rutledge & winkler (1989) made estimates of attenuation from vsp data in the upper basalt series in the vøring plateau area of the eastern norwegian sea. from 451 to 1111 m below the sea-floor in 1289 m of water, they found 105 basalt flows with about 10% of the section comprising volcaniclastic sediments. their estimates of overall scattering attenuation of 2.7 × 10–4 db/m (effective q = 25) could be accounted for by scattering loss modelled from the sonic and density logs recorded in the well, leaving an intrinsic attenuation of less than 0.6 × 10–4 db/m (q > 115). their q of 25 is somewhat less than the estimated effective q of 35 from the lopra-1/1a vsp. pujol & smithson (1991) reported values of effective q around 48, estimated from vsp data in thick, columbia plateau basalt sequences containing some interbedded clay zones. they also reported that scattering was the dominant loss mechanism, since elastic modelling was able to account for all the observed loss. the intrinsic losses in fig. 12. full elastic model of vsp downwave propagation through a stack of 3 m layers derived from the backus average of the vp, vs and density logs between 185 m below ground level (gl) and 2160 m below gl. propagation was modelled with and without peg-leg multiples and amplitudes were estimated in a 40 ms window about the wavelets. * 0 600 one-way loss (geometrical spreading and transmission losses with multiples) = 37.5 db one-way loss (geometrical spreading and transmission losses without multiples) = 46.9 db 6 x 10–3 4 x 10–3 2 x 10–3 –2 x 10–3 –4 x 10–3 –6 x 10–3 0 185 m rms amplitude windows wavelet without internal multiples with internal multiples 0 100 200 300 400 500 600 2160 m point source velocity model recorded time (ms) 6 x 10–5 4 x 10–5 2 x 10–5 –2 x 10–5 –4 x 10–5 –6 x 10–5 0 geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1937 38 basalt were low and could not be determined from the field data. the values of effective q estimated from lopra-1/1a is bracketed by those reported in the literature, and the general trends and inferred loss mechanisms are consistent. it is almost certain that more data points from west of the shetlands will result in a greater scatter of effective q estimates, corresponding to the variety of basalt environments. modelled offset vsp figure 14 shows the wide-band pand s-downwaves simulated at a horizontal array of receivers spaced 100 m apart and located 300 m below the base basalt in a half-space of sand. the volume injection source is located in a uniform, elastic layer, which is the overburden above the modelled basalt interval. the horizontal array of receivers in a 1d earth simulates a walkaway vsp shot into a single level geophone. multiples are included in the simulation. the direct p arrival is evident and its amplitude decays rapidly with offset due probably to a combination of strong geometrical spreading and critical angle effects at larger offsets in a medium with strong velocity contrasts. also evident is a strong, low-frequency event which dominates the shear record and is probably a mode conversion, propagating through the basalt sequence and emerging into the half space below. at the base basalt– sand contact, it converts to a p-wave and is also recorded as a strong event on the p-section, with an earlier arrival time. both events are visible over the offset interval 2000– 4500 m but neither event can be traced to zero offset. the asymptotic velocity of about 3.4 km/s is high and although it could correspond to the shear velocity of the shallow dolerites, which display the highest interval velocities in the sequence (fig. 2), the event arrives before the direct shear arrival curve and it must therefore have a compressional leg for part of its ray path. the limited offset interval makes interpretation difficult and such conversions may be sensitive to the particular velocity-depth function, but the observation offers some encouragement fig. 13: a: input and escaping wave spectra for a vsp downwave in a model similar to that in fig. 12, but with a half space representing a sand unit at the base of the basalt sequence. the source wavelet is white over the full spectrum to nyquist. the effects of scattering on the escaping wavelet are evident in the loss and the sidelobes. b: regression of spectral ratios of the downgoing wave referenced to the input wavelet. the estimated effective q is 32. 5.0e–4 6e–3 4.0e–4 3.0e–4 2.0e–4 1.0e–4 0.0e+0 4e–3 2e–3 0e–0 frequency (hz) 0 12010080604020 down-wave at 185 m down-wave at 2160 m a m pl itu de d ow nw av e at 1 85 m a m pl itu de d ow nw av e at 2 16 0 m 7e–3 5e–3 3e–3 1e–3 a –60 –50 –40 –30 –20 –10 0 regression estimate q = 32 frequency (hz) 0 90 80 70 60 50 40 30 20 10 sp ec tr al r at io ( db ) b least squares fit spectral ratios geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1938 39 for the use of mode-converted shear waves to probe below piled basalt flows, as described by emsley et al. (1998). conclusions the borehole seismic data recorded in the lopra-1/1a well show that the average p-wave velocity is high at 5248 m/s from mean sea level to the deepest geophone level at 3510 m kb. amplitude loss over the stacked basalt flows is moderate, corresponding to an effective q of 35. however, this amplitude loss can be modelled by geometrical spreading and elastic scattering, implying that intrinsic attenuation is low. persistent up-going events are evident within the interval logged by the vsp, even before wavefield separation, suggesting that lateral continuity of the basalt flow contacts is consistent over the several hundreds of metres that correspond to the vsp fresnel zone radius. we observed an unambiguous tie between the vsp and a primaries synthetic seismogram which allows a detailed correlation between the stratigraphy revealed by the drillbit with the events on the vsp reflected wavefield. both the pre-deepening targets at 920 ms and 1350 ms appear on the vsp at earlier times of 894 ms and 1330 ms. both events result from impedance contrasts within the volcanic sequence and it is likely that a strong reflection event, visible at 1412 ms twt is also within the volcanic sequence. its polarity indicates a decrease in acoustic impedance with depth so it is therefore unlikely to be basement. the prognosed depth of this event is 3732 m kb, or 167 m below the final td of the well at 3565 m kb. the unprocessed horizontal components suggest the presence of a persistent down-going shear wave, directly generated by the twin vibrators used as a vsp source, which in turn gives rise to up-going shear reflections. the vp/vs ratio from the vsp is 1.8 ± 0.1 (estimated error), which is in good agreement with the estimate of 1.84 ± 0.01 (one standard deviation) obtained from the pand s-sonic log data, and is rather constant over the logged interval. we estimate layer-induced anisotropy of about 5% due to the high elastic parameter contrasts in the basalt flows. azimuthal anisotropy estimated from the dipole shear log is low, but the direction of the anisotropy is consistent with mapped master joint sets: the well-defined ne–sw direction gives way to the less well-defined nw–se direction at about 540 m kb. our preferred explanation for the vertical p-wavespeed being 10% higher than the horizontal, as reported by kiørboe & petersen (1995), is that vertical velocities are locally raised by the presence of fast, dolerite intrusions. reflection seismic data are difficult to process and interpret in basalt covered areas. the lopra-1/1a borehole dataset offers insight into the seismic properties of basalts which we anticipate will be of benefit in designing and processing reflection surveys, a topic which has attracted considerable interest but brings with it acknowledged challenges. an immediate result is that ‘basalt’, in seismic terms, cannot be represented by a uniform slab of hard rock a couple of kilometres thick. we observe challenges in the geometrical spreading, scattering losses, multiple development and spectral colouring which point towards low frequencies as the best hope for imaging beneath bafig. 14. full elastic model of single level walkaway vsp with the geophone located 300 m below the basalt-sand contact. the left panel are p-waves while the right panel are s-waves. the ray-traced first p and pure s arrivals are superimposed. a mixed mode conversion can be seen with significant but low frequency amplitude over a limited offset interval. p s 400 800 1200 1600 2000 2200 50 45 50 45 50 50 50 offset (m) tw ow ay t im e (m se c) geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1939 40 salts. this observation has been reported elsewhere in the literature but this paper may allow some more specific numbers to be applied to the basalt flows in the faroese area. on the positive side, intrinsic attenuation is low and propagation has been demonstrated through 3533 m of basalts, tuffs and volcaniclastic sequences. coherent reflections have been tied from the vsp to the synthetic seismogram with confidence and another event below td has been prognosed. modelled offset vsp propagation also gave some hope for mode conversions, though with a limited offset range and low frequencies. acknowledgements this work was carried out while the first author was engaged on a secondment with bp in aberdeen. during this time he enjoyed considerable support from colleagues in bp, shell and schlumberger. in particular he would like to acknowledge many discussions and guidance from matt luheshi, cameron crook and brian mitchener. chris chapman answered many modelling questions and fraser louden recorded the vsp. references backus, g.e. 1962: long-wave elastic anisotropy produced by horizontal layering. journal of geophysical research 67, 4427–4440. dansk olieog gasproduktion & dansk operatørselskab 1997: end of well report, well lopra-1, 80 pp. + 11 appendices. thorshavn, faroe islands: jardfeingi/faroese earth and energy directorate. emsley, d., boswell, p. & davis, p. 1998: sub-basalt imaging using long-offset reflection seismic data. 60th meeting, european association of geoscientists and engineers expanded abstract, 1– 48. esmersoy, c., koster, k., williams, m., boyd, a. & kane, m. 1994: dipole shear anisotropy logging. 64th annual international meeting, society of exploration geophysicists, expanded abstracts 1139–1142. folstad, p.g. & schoenberg, m. 1992: low frequency propagation through fine-layering. 62nd annual international meeting, society of exploration geophysicists, expanded abstracts 1278– 1281. folstad, p.g. & schoenberg, m. 1993: scattering from a set of anisotropic layers to second order in frequency. 55th meeting, european association of exploration geophysicists, extended abstracts paper, 105 only. kennett, b.l.n. 1974: reflections, rays and reverberations. bulletin of the seismological society of america 64, 1685–1696. kennett, b.l.n. 1983: seismic wave propagation in stratified media, 342 pp. cambridge: cambridge university press. kiørboe, l. & petersen, s.a. 1995: seismic investigation of the faroe basalts and their substratum. in: scrutton, r.a. et al. (eds): the tectonics, sedimentation and palaeoceanography of the north atlantic region. geological society special publication (london) 90, 111–123. larsen, l.m., waagstein, r., pedersen, a.k. & storey, m. 1999: trans-atlantic correlation of the palaeogene volcanic successions in the faeroe islands and east greenland. journal of the geological society (london) 156, 1081–1095. newman, p. 1973: divergence effects in a layered earth. geophysics 38, 481–488. nielsen, p.h., stefánsson, v. & tulinius, h. 1984: geophysical logs from lopra-1 and vestmanna-1. in: berthelsen, o., noenygaard, a. & rasmussen, j. (eds): the deep drilling project 1980–1981 in the faroe islands, 115–135.tórshavn: føroya fróðskaparfelag. o’doherty, r.f. & anstey, n.a. 1971: reflections on amplitudes. geophysical prospecting 19, 430–458. planke, s. & cambray, h. 1998: seismic properties of flood basalts from hole 917a downhole data, southeast greenland volcanic margin. in: sanders, a.d., larsen, h.c. & wise, s.w. jr. (eds): proceedings of the ocean drilling program, scientific results 152, 453–462. pujol, j. & smithson, s.b. 1991: seismic wave attenuation in volcanic rocks from vsp experiments. geophysics 56, 1441–1455. rasmussen, j. & noe-nygaard, a. 1970: geology of the faeroe islands. danmarks geologiske undersøgelse 1. række 25, 142 pp. rasmussen, j. & noe-nygaard, a. 1990: the origin of the faeroe islands in text, pictures and on maps, 64 pp., 6 maps at 1:50 000. copenhagen: geological survey of denmark (also text in faeroese and danish). rutledge, j.t. & winkler, h. 1989: attenuation measurements from vertical seismic profile data: leg 104, site 642. in: eldholm, o. et al. (eds): proceedings of the ocean drilling program, scientific results 104, 965–972. thomsen, l. 1986: weak elastic anisotropy. geophysics 51, 1954– 1966. white, r. & mckenzie, d. 1989: magmatism at rift zones: the generation of volcanic continental margins and flood basalts. journal of geophysical research 94, 7685–7729. manuscript received 16 october 2001; revision accepted 14 march 2003. geus bulletin no 9 7 juli.pmd 07-07-2006, 14:1940 geological survey of denmark and greenland bulletin 17, 2009 carbon capture and storage (ccs) is increasingly considered to be a tool that can significantly reduce the emission of co2. it is viewed as a technology that can contribute to a substantial, global reduction of emitted co2 within the timeframe that seems available for mitigating the effects of present and continued emission. in order to develop the ccs method the european union (eu) has supported research programmes for more than a decade, which focus on capture techniques, transport and geological storage. the results of the numerous research projects on geological storage are summarised in a comprehensive best practice manual outlining guidelines for storage in saline aquifers (chadwick et al. 2008). a detailed directive for geological storage is under implementation (european commission 2009), and the eu has furthermore established a programme for supporting the development of more than ten large-scale demonstration plants throughout europe. geological investigations show that suitable storage sites are present in most european countries. in denmark initial investigations conducted by the geological survey of denmark and greenland and private companies indicate that there is significant storage potential at several locations in the subsurface. the danish perspective in storage capacity the ten largest point sources of co2 emission in denmark account for 21 mega-tonnes per year (mt/year). from preliminary investigations of the danish subsurface the co2 storage capacity in selected subsurface structures is estimated to 2500 mt (geocapacity 2009a). this corresponds to more than 100 years of storage from the ten largest emission point sources. the critical parameters of this analysis are the size of the structure, thickness, continuity and quality of the reservoir and the amount of formation water that may be displaced by the injected co2. the estimate is calculated assuming a surrounding aquifer volume displacement of formation water, which is limited to 50 times the trap volume (geocapacity 2009b). these estimates of storage capacity are uncertain and have not yet been tested in real physical storage operation. therefore it is difficult to evaluate to what degree the volume calculations are realistic. when a specific structure is selected for storage, a number of investigative steps are necessary, including acquisition and interpretation of new 2-d or 3-d seismic data, drilling of new wells, geological and reservoir modelling and flow simulation studies. for each step of incorporating new geological data, the site model of the reservoir is updated in the process of maturing the structure towards a storage site. this stepwise approach to site characterisation gradually leads to a research-based and relatively certain capacity estimate and an evaluation of the safety and behaviour of the site under simulated conditions, including the uncertainties of the estimates. assessment of geological and environmental risks can be carried out at various stages in the process. similarly, establishment of baseline studies and monitoring strategies need to be considered along with the progress of the characterisa© geus, 2009. geological survey of denmark and greenland bulletin 17, 13–16. available at: www.geus.dk/publications/bull the potential for large-scale, subsurface geological co2 storage in denmark peter frykman, lars henrik nielsen, thomas vangkilde-pedersen and karen lyng anthonsen 13 �� �� �� �� �� �� �� �� �� �� vedsted structure stenlille structure jylland fyn sjælland ålborg large point sources structures �� 50 km fig. 1. map of denmark showing the most important point sources of co2 emission and prospective structures for geological storage of co2. the stenlille structure is presently used for storage of natural gas; it serves to moderate seasonal fluctuations in consumption. the vedsted structure is currently investigated for possible storage of co2. rosa_2008:rosa-2008 01/07/09 15:47 side 13 tion of the storage site in order to make sure that the necessary background information is obtained before storage is initiated. in order to be reliable and operational, the baseline studies should preferably focus on measurement of conditions and properties that are stable and only show limited seasonal variations. such studies may include groundwater-flow models and groundwater chemistry, pore water and pore gas analyses from deep wells, surface topography and natural seismicity. a site study site investigations have recently been initiated of the vedsted structure by vattenfall a/s, with the intention of using the structure for storage of co2 from a nearby coal-fired power plant in ålborg (fig. 1; sørensen et al. 2009). existing data from oil exploration activities in the 1950s include one well in the centre of the structure and sparse 2-d seismic line data. the main target layer is the triassic–jurassic gassum for mation at around 1800 m depth. the formation is widely distributed in the danish basin and has good reservoir properties (fig. 2). it is currently used for storage of natural gas in the stenlille structure on sjælland and for geothermal energy in the thisted area in northern jylland. detailed sedimentological and sequence stratigraphic interpretations and correlations of the well logs and cores have established a robust stratigraphic framework for the upper triassic – jurassic succession (nielsen 2003). this framework forms the basis for the interpretation of the vedsted-1 well section as well as predictions regarding the lithology of the potential reservoirs and seals in the vedsted area (fig. 3). the process has also underlined the necessity of acquisition of new data and more detailed modelling at several different scales. at site scale, the optimal positioning of injection wells, as well as injectivity and capacity can be modelled and analysed, and the coupling between the operation of the power plant and the capture facility can be studied. the specific geological properties of the storage reservoir layers have consequences for the propagation and distribution of the injected co2 and for the storage mechanisms in the specific reservoir. most reservoirs show both vertical and horizontal heterogeneities that will influence the distribution of the co2. the preliminary reservoir model for the vedsted structure has been investigated by simulating an injection well on its south-eastern flank and using injection rates realistic for power-plant supply rates (frykman et al. 2009). after ten years of constant injection, the co2 distribution is as seen in fig. 4, which clearly shows the subdivision of the migrating front into several sub-layers due to intraformational sealing layers with low permeability that also have high capillary entry pressures. the layering in the model has maximum lateral continuity, which probably overestimates the segregation to be found in real cases, but any intra-reservoir sealing layers will have such an effect on the distribution. since this filling pattern influences the capacity, it is necessary to analyse further the properties and the continuity of the intraformational sealing layers. co2 can be trapped by several mechanisms, including structural trapping under an overlying sealing formation, dissolution of co2 in formation water, capillary trapping in the pore network and mineral trapping by reactions between co2 and mineral phases in the reservoir rock. these trapping mechanisms work on different scales both in space and time and need to be studied by designing appropriate models and experiments. for large-scale injection of co2 displacing saline porewater, the propagation of the pressure field during injection outside the immediate site area is of interest. modelling of this pressure distribution will serve to predict the amount of overpressure building up locally within the storage site, and can be used to suggest possible means of management. 14 50 km fig. 2. distribution of the triassic–jurassic gassum formation in the subsurface at depths between 800 m and 2400 m (yellow), the depth interval in which co2 exists as a supercritical phase and where burial dia genesis has not yet provoked significantly lowered porosity. at the supercritical phase the volume of co2 is much less than that of the co2 gas at the surface. rosa_2008:rosa-2008 01/07/09 15:47 side 14 the scale of the challenge and future perspective the first detailed pan-european assessment of co2 storage capacity in the framework of the eu research project geo capacity has resulted in a geographic information system (gis) database of co2 emissions, storage capacity estimates and geological information. the database includes information on reservoirs with a total storage capacity of 360 000 mt co2, with 326 000 mt in deep saline aquifers, 32 000 mt in depleted hydrocarbon fields and 2000 mt in unmineable coal beds: 116 000 mt are onshore, and 244 000 mt offshore (geocapacity 2009a). some of the estimated storage capacity is associated with structural traps, but a very large part is in regional deep saline aquifers without identified specific traps. almost 200 000 mt of the total storage capacity in the database are located offshore norway. these estimates date back to 2003 and have not been updated within the geocapacity project. an attempt to provide a more cautious and conservative european estimate has yielded a storage ca pacity of 117 000 mt with 96 000 mt in deep saline aquifers, 20 000 mt in depleted hydrocarbon fields and 1000 mt in coal beds, and with approximately 25% located offshore norway. this must be compared to a total of 2000 mt of co2 emission from large point sources, i.e. point sources emitting more than 0.1 mt/year within europe. in order to illustrate the scale of the technology and infrastructure that has to be established if ccs is to become an active industry, we can look at the amount of co2 produced by the ten largest point sources in denmark. there are 43 large point sources emitting 28 mt co2/year, the ten largest of which are responsible for 21 mt/year. at surface condi15 fig. 3. sw–ne-oriented cross-section across the danish basin, the sorgenfrei–tornquist zone and the skagerrak–kattegat platform (red line on the index map). the panel shows the lower part of the gassum reservoir which comprises fluvial, estuarine and shallow-marine deposits interbedded with offshore mudstones and some lacustrine mudstones. shoreline fluctuations have caused interfingering of these different facies types and given rise to pronounced vertical variability. informations from the four wells in the section about sedimentary facies have been interpreted and correlated into a sequence-stratigraphic framework. at a local site, this framework must be confirmed from detailed investigations of material from new wells drilled, and supplemented with new seismic data. modified from nielsen (2003). 4 ce rl rl skagerrak–kattegat platform sorgenfrei–tornquist zone 37 km 5 km 41 km 30 km 20 km børglum fault ation pt pt ts 7 danish basin sæby-1 vedsted-1rødding-1 sp hyllebjerg-1farsø-1 gr flyvbjerg-1 nesw 32 50 m 50 km sorgenfrei–tornquist zone danish basin skagerrak–kattegat platform fyn high ringkøbing – fluvial estuarine lacustrine lagoonal shoreface offshore depositional environments rosa_2008:rosa-2008 01/07/09 15:47 side 15 tions this corresponds to 11 billion (11 × 109) m3 co2 gas. the annual production of natural gas from the danish part of the north sea amounts to 10 billion (10 × 109) m3 (danish energy agency 2008), which is transported in pipe lines and tankers and processed at plants and refineries. the comparable size of the potential volume of co2, to be moved around at surface and injected into the subsurface (although compressed to smaller volumes at depth), points to the large scale at which a ccs-related processing and transporting industry has to be established. concluding remarks the ccs activities described here related to large-scale storage operations will involve significant physical resources and manpower. fortunately, the work with storage-related items does not have to begin from first principles, because much of the experience already exists in the oil and gas industry, which can provide methods and tools for immediate use. the skills of geoscientists and engineers are needed in the investigation and characterisation of the sites and the subsurface conditions for storage of co2, and a whole new infrastructure and industry may be established. geoscience and geo-engineering will play a major role in the analysis of the geological foundation, the assessment of site performance, and will be critical in securing the safety of the operations. initial investigations of the danish subsurface indicate that suitable structural traps with a significant storage potential are present at several locations, and that the structures can accommodate the co2 produced from several or most of the large danish point sources. thus, geological storage of co2 may contribute considerably to the reduction of the danish co2 emission, if we can be assured about safety issues, and if political and public acceptance can be obtained. references chadwick, a., arts, r., bernstone, c., may, f., thibeau, s. & zweigel, p. 2008: best practice for the storage of co2 in saline aquifers – observations and guidelines from the sacs and co2store projects. british geological survey occasional publication 14, 267 pp. danish energy agency 2008: oil and gas production in denmark 2007, 98 pp. copenhagen: danish energy agency. european commission 2009: directive of the european parliament and of the council on the geological storage of carbon dioxide, 84 pp. brussels: european union. frykman, p., bech, n., sørensen, a.t., nielsen, l.h., nielsen, c.m., kristensen, l. & bidstrup, t. 2009: geological modelling and dynamic flow analysis as initial site investigation for large-scale co2 injection at the vedsted structure, nw denmark. 9th international conference on greenhouse gas control technologies, washington d.c., 16–20 no vember, 2008. ghgt9 energy procedia 1, 2975–2982. geocapacity 2009a: geocapacity wp 2 report, storage capacity. eu geocapacity deliverable d16, 162 pp. eu geocapacity consortium, brussels. geocapacity 2009b: geocapacity wp 4 report, capacity standards and site selection criteria. eu geocapacity deliverable d26, 45 pp. eu geocapacity consortium, brussels. nielsen, l.h. 2003: late triassic – jurassic development of the danish basin and the fennoscandian border zone, southern scandinavia. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and green land. geological survey of denmark and greenland bulletin 1, 459–526. sørensen, a.t., klinkby, l., christensen, n.p., dalhoff, f., biede, o. & noer. m. 2009: danish development of a full-scale ccs demonstration plant in a saline aquifer. first break 27, 79–83. 16 authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: pfr@geus.dk fig. 4. vertical nw–se section in the gassum reservoir model through the injection well, showing co2 saturation sg (free gas-phase supercritical co2) after 10 years of injection. although the model is constructed in a fairly coarse grid, the intra-reservoir sealing layers are clearly reflected and influence the spatial distribution of the injected co2. the sedimentary layering causes filling of the individual layers of porous sand with the injected co2, whereas the interbedded mudstone layers with much lower reservoir quality are not filled and also limit the vertical movement of co2. model thickness 300 m, length 4800 m, vertical exaggeration 5 times. (modified from frykman et al. 2008). 0.55 sg 0.5 0.45 0.4 0.35 0.3 0.25 0.2 0.15 0.1 0.05 0 rosa_2008:rosa-2008 01/07/09 15:47 side 16 geological survey of denmark and greenland. bulletin 10, 25-28 laser ablation inductively coupled plasma mass spectrometry (la-icp-ms) was developed in 1985 and the first commercial laser ablation systems were introduced in the mid 1990s. since then, la-icp-ms has become an important analytical tool in the earth sciences. initially, the main interest for geologists was in its ability to quantitatively determine the contents of a wide range of elements in many minerals at very low concentrations (a few ppm and below) with relatively high spatial resolution (spot diameters of typically 30–100 μm). the potential of la-icp-ms for rapid in situ u–th–pb geochronology was already realised in the early to mid 1990s. however, the full potential of la-icp-ms as the low-cost alternative to ion-microprobe techniques for highly precise and accurate in situ u–th–pb age dating was not realised until the relatively recent advances in laser technologies and the introduction of magnetic sectorfield icp-ms (sf-icpms) instruments. in march 2005, the geological survey of denmark and greenland (geus) commissioned a new laser ablation magnetic sectorfield inductively coupled plasma mass spectrometry (la-sf-icp-ms) facility employing a thermofinnigan element2 high resolution magnetic sectorfield icp-ms and a merchantek new wave 213 nm uv laser ablation system. the new geus la-sf-icp-ms facility is widely used on survey research projects in denmark and greenland, as well as in collaborative research and contract projects conducted with partners from academia and industry worldwide. here, we present examples from some of the these ongoing studies that highlight the application of the new facility for advanced geochronological and trace element in situ microanalysis of geomaterials. the application of lasf-icp-ms based in situ zircon geochronology to regional studies addressing the archaean geology of southern west greenland is presented by hollis et al. (2006, this volume). zircon u–pb geochronology using la-sf-icp-ms in situ u–th–pb geochronology was developed in the mid80s with the introduction of ion-microprobe techniques, most commonly referred to as secondary ion mass spectrometry (sims) and sensitive high resolution ion microprobe (shrimp). the advantage of in situ u–th–pb geochronology over conventional chemical dating by isotope dilution thermal ionisation mass spectrometry (id-tims) is the capability to analyse different domains in heterogeneous single zircons with high spatial resolution (spot diameters of typically 10–30 μm). this allows resolution of igneous and metamorphic events separated by intervals of only a few tens of million years from polychronic zircons. the disadvantages of ion-microprobe techniques are the very high purchasing and operating costs for the instrument. the rapid improvements in laser based u–th–pb geochronology makes it now possible to obtain in situ u–th–pb geochronological data with comparable spatial resolution as well as analytical precision and accuracy at only a fraction of the costs of ion-microprobe techniques (e.g. jackson et al. 2004; janoušek et al. 2006). dating of magmatic and metamorphic events the capabilities of la-sf-icp-ms for the precise and accurate u–pb age dating of relatively young igneous zircons are demonstrated by the analysis of a population of 40 zircons extracted from a gabbro from the coastal cordillera at tregualemu, central southern chile. the gabbro is believed to have been formed by regional extension during the late triassic to early jurassic (charrier 1979). three zircons proved to be too small for analysis (< 30 μm). the results for the remaining 37 zircons (fig. 1) define a highly precise igneous concordia age of 203 ± 2 ma (2σ; mswd = 1.7) and indicate a late triassic (rhaetian) intrusion age of the gabbro. an even advanced in situ geochronological and trace element microanalysis by laser ablation techniques dirk frei, julie a. hollis,axel gerdes, dan harlov, christine karlsson, paulina vasquez, gerhard franz, leif johansson and christian knudsen © geus, 2006. geological survey of denmark and greenland bulletin 10, 25–28. available at: www.geus.dk/publications/bull 25 240 220 200 180 1600.025 0.027 0.029 0.031 0.033 0.035 0.037 0.17 0.19 0.21 0.23 0.25 0.27 207 pb/ 235 u 2 06 p b /2 38 u geus la-sf-icp-ms: igneous concordia age 203 ± 2 ma (95% conf.) mswd: 1.7 gabbro, central southern chile (sample pv 4-22) fig. 1. concordia diagram for igneous zircons from a gabbro from the coastal cordillera in central southern chile. younger igneous age of 158 ± 2 ma was recently obtained for a zircon xenoor phenocryst derived from a newly discovered carbonatite in southern west greenland (steenfelt et al. 2006, this volume). the high sensitivity of the element2 sf-icp-ms allows u–pb zircon age dating with a laser spotsize of 30 μm or less, depending on the pb content of the zircons. this makes it feasible to analyse different age domains in polychronic zircons, e.g. igneous cores and metamorphic rims. for example, zircons from an orthogneiss from the nuuk region, southern west greenland, display characteristic textures in back-scattered electron (bse) and cathodoluminescence (cl) pictures that are interpreted as igneous cores surrounded by rims grown during a metamorphic event (see inset in fig. 2). the u–pb age data of the cores suggest an emplacement of the igneous protolith at c. 3660 ma, while the rim data indicate metamorphism close to 2700 ma (fig. 2). dating of detrital zircons analyses of the crystallisation ages of detrital zircons in clastic sediments are a powerful tool in sedimentary provenance analysis. accurate and precise u–pb ages of >100 detrital zircon grains in a sample are needed to detect all major sedimentary source components with statistical confidence (cf. vermeesch 2004; and references therein). the relatively high costs and the limited capacities of ion microprobe techniques (c. 75 zircon age analyses per day) impose restrictions on the number of samples that can be studied. because la-sf-icpms provides very high capacities (in excess of 300 zircon age analyses per day) without compromising accuracy and precision, it constitutes the economic method of choice for provenance studies based on detrital zircon u–pb ages. an example for detrital zircon age data obtained by lasf-icp-ms is shown in fig. 3, where the 207pb–206pb age distribution for a population of 100 zircon grains separated from a cambrian sandstone from torekov, southern sweden, are shown in a combined histogram and probability density distribution (ppd) diagram. the concordance filtered zircons (dark shaded area; 90–110% concordance, defined as 100* [206pb–238u age / 207pb–206pb age]) show a polymodal age distribution with a minor peak at ~1000 ma and two major peaks at c. 1150 ma and 1650 ma. the presence of two older sedimentary sources (c. 2150 ma and c. 3050 ma) is indicated by discordant grains (lighter shaded grey areas) that most likely suffered lead loss during their petrogenetic evolution. figures of merit the short and long term precision and accuracy of la-sficp-ms for u–pb zircon age dating has been assessed using two zircon reference materials, plesovice (with an id-tims age of 338 ± 1 ma; aftalion et al. 1989; provided by jan kosler, university of bergen) and 91500 (id-tims age = 1065 ± 0.4 ma; wiedenbeck et al. 1995). the pl zircon is routinely analysed as unknown for quality control purposes in every analytical session in the geus laboratory. the results for 16 analyses of the zircon from a typical single analytical session are shown in fig. 4a. they define a concordia age that is in excellent agreement with the id-tims age reported by aftalion et al. (1989). long-term precision (2σ) based on 109 analyses of the plesovice zircon by two different operators was 2%, 2.3% and 1.1% for the 206pb/238u, 207pb/235u and 207pb/206pb ratios, respectively. the widely used 91500 zircon has so far only been analysed during one analytical session. the results for all seven analyses carried out 26 3900 3700 3500 3300 3100 2900 2700 2500 3800 3400 3000 2600 2200 c. 2700 ma n=8 c. 3660 ma n=4 2 0 6 p b /2 3 8 u 207 pb/ 235 u 0.9 0.8 0.7 0.6 0.5 0.4 0.3 5 15 25 35 45 fig. 2. concordia diagram and inset of 207pb/206pb age plot for polychronic zircons from an orthogneiss in the nuuk region, southern west greenland. inset cl image shows laser ablation pits (pit diameter = 30 μm) in igneous cores and metamorphic rims of polychronic zircons. note the shallow depth (usually < 20 μm) of the laser ablation pits. 0 40 0 8 00 1 20 0 16 00 2 00 0 2 40 0 28 00 3 20 0 3 60 0 40 00 age (ma) p ro b ab ili ty 0 5 10 15 20 25 30 35 f re q u en cy fig. 3. combined display of histogram and probability density distribution (ppd) diagram for zircons from a cambrian sandstone from torekov, south-western sweden. see text for explanations. during this session (fig. 4b) define a concordia age which is in excellent agreement with the id-tims age reported by wiedenbeck et al. (1995). trace element analysis using la-sf-icp-ms due mainly to the introduction of la-icp-ms techniques, the trace element signatures of individual minerals (e.g. garnet, clinopyroxene, epidote, rutile, calcite) are now frequently being used to deduce the petrogenetic evolution of magmatic rocks, unravel water–rock interactions, identify geotectonic settings and sediment sources, track down the pathways of potentially health-damaging pollutants, and to unravel the change of seasurface and atmospheric temperatures. furthermore, la-icp-ms is an important tool for the characterisation of synthetically produced geomaterials. water–rock interaction fluid-mediated mass transfer during metamorphism and metasomatism is a hotly debated issue. mobility of geochemically important trace elements during fluid flow has farreaching implications for e.g. ore-forming processes and mass transfer during subduction. detailed investigation of the mechanisms of element mobility during fluid–rock interaction on a grain-scale are pivotal for an understanding of the processes that lead to the characteristic element enrichments and depletions observed in nature. in the söndrum stone quarry in sweden, a localised dehydration zone of 2.5 to 3 m width occurs in garnet-bearing granitic gneiss around an approximately 1 m wide pegmatoid dyke. whole-rock chemistry suggests that the solid-state dehydration of the granitic gneiss to charnokite via low h2o activity fluids consisting principally of co2 and a minor brine component was predominantly isochemical (harlov et al. 2006). exceptions include y and the heavy rare earth elements (hree), which are markedly depleted throughout the dehydration zone. in order to assess the mechanism of y and hree depletion, the trace element geochemistry of garnet was studied in a traverse across the dehydration zone. garnets from the pristine, unaltered granitic gneiss are characterised by a strong negative eu-anomaly and the steep, hree enriched pattern typical for garnet (sample sd45-600 in fig. 5). in contrast, garnets from within the dehydration zone show a less pronounced eu-anomaly and are characterised by dramatic y and hree depletions, leading to almost flat ree pattern (sample sd9-120 in fig. 5). this observation provides direct evidence for massive release of y and hree from garnets, the principal hosts of these elements in the granitic gneiss, via solid-state fluid–rock interaction which was also accompanied by dehydration of hornblende and biotite to orthoand clinopyroxene. analysis of synthetic geomaterials another application of la-icp-ms is the trace element analysis of synthetic geomaterials, e.g. products of experiments carried out for petrogenetical purposes. since the experimentally produced mineral phases are usually very small, in situ microanalysis with high spatial resolution is needed. this was traditionally achieved by sims techniques. because laicp-ms analyses are much cheaper and facilities are much 380 360 340 320 300 0.047 0.049 0.051 0.053 0.055 0.057 0.059 0.061 0.34 0.36 0.38 0.40 0.42 0.44 0.46 geus la-sf-icp-ms: concordia age 339.0 ± 2.7 ma (95% conf.) mswd: 2.1 a: plesovice zircon standard id-tims age: 338 ± 1 ma 980 1020 1060 1100 1140 1180 0.16 0.17 0.18 0.19 0.20 1.6 1.7 1.8 1.9 2.0 2.1 2.2 geus la-sf-icp-ms concordia age 1070 ± 10 ma (95% conf.) mswd: 0.27 b: 91500 zircon standard id-tims age: 1065 ± 0.4 ma 207 pb/ 235 u 207 pb/ 235 u 2 06 p b /2 38 u 2 06 p b /2 38 u fig. 4. concordia diagrams for 16 analysis of the plesovice zircon reference material obtained during a single analytical session (a); and 7 analysis of the 91500 zircon reference material obtained during a single analytical session (b). 27 sm eu gd tb dy ho er tm yb lu c 1 n o rm al is ed sd45-600 sd9-120 1 10 100 1000 10000 fig. 5. average chondrite-normalised ree-pattern of garnets from the unaffected granitic gneiss (sd45-600) and the dehydration zone (sd9120). note the depletion in hree of garnets from the dehydration zone. more widely available, la-icp-ms is increasingly used for the trace element analysis of synthetic geomaterials. the inset in fig. 6, for example, shows a back-scattered electron photomicrograph of an experimental charge containing euhedral clinopyroxene and a coexisting anhydrous silicate melt. the experimental charge was synthesised in order to determine the clinopyroxene-melt trace element partition coefficients di (defined as di = [concentration of i]cpx/[concentration of i]melt), knowledge of which is important for geochemical melt modelling purposes (landwehr et al. 2001). in order to test the ability of laser ablation techniques to correctly determine partition coefficients from such experimental charges, we have analysed clinopyroxenes and coexisting silicate melts from a number of experiments with both sims (at the nerc ion-microprobe facility in edinburgh) and la-sf-icp-ms (at geus). the resulting partition coefficients determined by sims and la-sf-icp-ms are graphically compared in fig. 6 for a representative sample. for all investigated samples, the partition coefficients determined by sims and la-sf-icp-ms are in excellent agreement, clearly demonstrating the reliability of laser ablation techniques for the characterisation of synthetic geomaterials. acknowledgements the establishment of the new la-sf-icp-ms facility was funded with a grant of the danish ministry of education and technology to the geocenter copenhagen (‘geocenterbevilling’) and financial support by geus. pv thanks the daad for awarding a phd scholarship. references aftalion, m., bowes, d.r. & vrána, s. 1989: early carboniferous u–pb zircon age of garnetiferous, perpotassic granulites, blanský les massif, czechoslovakia. neues jahrbuch für mineralogie, monatsheft 4, 145–152. charrier, r. 1979: el triásico en chile y regiones adyacentes de argentina: una reconstrucción paleogeográfica y paleoclimática. comunicaciones 26, 1–37. harlov, d.e., johansson, l., van den kerkhof, a. & förster, h.-j. 2006: the role of fluid flow and diffusion during localised, solid-state dehydration: söndrum stenhuggeriet, halmstad, sw sweden. journal of petrology 47, 3–33. hollis, j.a., frei, d., van gool, j.a.m., garde, a.a. & persson, m. 2006: using zircon geochronology to resolve the achaean geology of southern west greenland. geological survey of denmark and greenland bulletin 10, 49–52. jackson, s., pearson, n.j., griffin, w.l. & belousova, e.a. 2004: the application of laser ablation – inductively coupled plasma – mass spectrometry to in situ u–pb zircon geochronology. chemical geology 211, 47–69. janoušek, v., gerdes, a., vrána, s., finger, f., erban, v., friedl, g. & braithwaite, c.j.r. 2006: low-pressure granulites of the lišov massif, southern bohemia: viséan metamorphism of late devonian plutonic arc rocks. journal of petrology 47, 705–744. landwehr, d., blundy, j.d., chamorro-perez, e.m., hill, e. & wood, b.j. 2001: u-series disequilibria generated by partial melting of spinel lherzolite. earth and planetary science letters 188, 329–348. steenfelt, a., hollis, j.a. & secher, k. 2006: the tikiusaaq carbonatite and associated kimberlites: a new alkaline magmatic province in the nuuk region, southern west greenland. geological survey of denmark and greenland bulletin 10, 41–44. wiedenbeck, m., allé, p., corfu, f., griffin, w.l., meier, m., oberli, f., von quadt, a., roddick, j.c. & spiegel, w. 1995: three natural zircon standards for u–th–pb, lu–hf, trace element and ree analysis. geostandards newsletters 19, 1–23. vermeesch, p. 2004: how many grains are needed for a provenance study? earth and planetary science letters 224, 441–451. authors’ addresses d.f., j.h. & c.kn., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: df@geus.dk a.g., institute of mineralogy, johan-wolfgang-goethe university, senckenberganlage 28, d-60054 frankfurt, germany. d.h., geoforschungszentrum potsdam, section 4.1 experimental geochemistry and mineral physics, telegrafenberg, d-14473 potsdam, germany. c.ka. & l.f., department of geology, university of lund, sölvegatan 12, s-22362 lund, sweden. p.v. & g.f., institut für angewandte geowissenschaften, technische universität berlin, ernst-reuter-platz 1, d-10587 berlin, germany. 28 10 -3 10 -2 10 -1 10 0 10 1 10 -3 10 -2 10 -1 10 0 10 1 la ce nd sm lu ti zr hf th u d si m s d la-sf-icp-ms 10 -3 10 -2 10 -1 10 0 10 1 10 -3 10 -2 10 -1 10 0 10 1 experimentally determined cpx-melt partition coefficients la ce nd sm lu ti zr hf th u d si m s d la-sf-icp-ms 500 μm fig. 6. comparison of clinopyroxene-melt trace element partition coefficients (open symbols: ree; solid symbols: hfse) determined from synthetic run products by sims and la-sf-icp-ms. inset shows a bse photomicrograph of synthetic, euhedral clinopyroxene coexisting with a glassy melt produced experimentally at high pressure and temperature. shallow, bright spots are sims ablation pits (c. 20 μm diameter), while deeper, dark spots are laser ablation pits (c. 30 μm in diameter). geological survey of denmark and greenland bulletin 23, 2011, 73–76 73 the greenland ice sheet is reacting to climate change. yet, mass-budget estimates differ considerably, partly due to climatic variability and partly to uncertainties in the techniques of assessing mass change (ipcc 2007). nevertheless, all recent estimates agree that the ice sheet is losing mass (e.g. 286 gt/yr; velicogna 2009) at an accelerating rate (rignot et al. 2011). on top of this, the area with a negative mass budget is expanding rapidly (khan et al. 2010). the mass loss is attributed equally to increases in both iceberg production and melting of the ice sheet (van den broeke et al. 2009). the increasing mass loss in recent years has caught public attention and given rise to concern worldwide due to its potential impact on sea level. in the light of this, the programme for monitoring of the greenland ice sheet (promice) was initiated in 2007 (ahlstrøm & promice project team 2008), lead by the geological survey of denmark and greenland (geus). promice undertakes surface mass-budget measurements using automatic weather stations, quantifies the mass loss by iceberg calving using remotely sensed data from satellites and airborne surveys and tracks changes in the extent of glaciers. in this paper, we focus on weather station measurements, which are crucial in calculating the energy exchange between the atmosphere and the ice sheet, and in validating model calculations of the surface mass budget. in particular, we present the observed temperatures and investigate how their high 2010 values affected ablation in southern greenland. promice automatic weather stations the promice weather station network started with five stations in 2007 and by summer 2010 consisted of seven station pairs (fig. 1; table 1). typically, one of the stations in a pair is located in the upper ablation zone near the equilibrium line and the other at a lower elevation well into the ablation zone. the weather stations are equipped with the instruments shown in fig. 2 which undergo continuous 10-minute measurement cycles. in summer, data are transmitted once per hour; in winter, transmissions are daily to reduce power consumption when solar power is limited. promice weather station data can be downloaded at no charge at www.promice.dk. in spite of the stations being placed in inhospitable places where strong winds, severe cold, icing as well as melting and highly uneven terrain are common, there was a success rate of 77–86% for the period up to february 2011. not all data have been transmitted with success, so the success rate may reach 86% when also locally programme for monitoring of the greenland ice sheet (promice): first temperature and ablation records dirk van as, robert s. fausto and the promice project team* *andreas p. ahlstrøm, signe b. andersen, morten l. andersen, michele citterio, karen edelvang, peter gravesen, horst machguth, faezeh m. nick, søren nielsen and anker weidick. 80°n 75°n 70°n 65°n 60°n 80°w 60°w 40°w 50°w 40°w 20°w 20°w 0°w 2250 3000 2750 2500 1000 1500 1750 20 0022 50 27 5025 00 22 50 17 50 1500 1 2 5 0 2000 greenland 500 km qas nuk upe thu kpc sco tas fig. 1. map of greenland with the locations of the promice automatic weather stations in 2010. each dot represents a pair of stations. station abbreviations as in table 1. dotted lines: elevation contours. © geus, 2011. geological survey of denmark and greenland bulletin 23, 73–76. open access: www.geus.dk/publications/bull 7474 stored data have been collected. strikingly, only few values are missing due to harsh climatic conditions such as wind damage. a prime cause of data gaps is data logger malfunction. temperatures over the ice sheet in fig. 3 we show the monthly mean near-surface air temperatures at those promice stations for which data cover at least half a month. a clear annual cycle is present in the temperature records of all weather stations, and we see that the amplitude of the annual signal increases with latitude. this is explained by the fact that during summer the solar radiation increases with latitude due to the midnight sun, while the opposite is true during winter when central and northern greenland experiences polar night. during the ‘warm’ season the presence of a melting ice surface at the stations does not allow near-surface temperatures to increase well above freezing. in southern greenland, where day-time, free atmospheric temperatures can exceed 20°c during summer, the melting ice surface dampens the amplitude of the temperature cycle by about 10°c. the smallest amplitude and highest winter temperatures occur at the tasiilaq stations (tas_l and tas_u), and the lower qassimiut station (qas_l; fig. 1; table 1). these stations are located at lower elevations close to the ice-sheet margin, and are exposed to the relatively warm wintertime atmospheric conditions of the atlantic ocean. the largest amplitude in the temperature cycle is seen at the upper kronprins christian land station (kpc_u), where melting occurs in summer, but where mean temperatures drop below –30°c in winter. the lowest daily mean temperature recorded at this station was –40.6°c on 9 january 2010. mid summer (july) monthly mean temperatures are above freezing at all stations, but never exceed 6°c. temperatures in greenland have been rising since the 1980s, prior to which there was almost half a century of cooling (box 2002). still, 2010 was exceptionally warm over large parts of greenland. it was the warmest year in greenland on record at most of the land-based weather stations operated by the danish meteorological institute. the only exception was seen in the north-east. individual months and seasons showed record setting temperatures, with the longest instrument records in greenland dating back to the 1870s (john cappelen, personal communication, 2011). the promice weather station network was not fully established until 2010 and thus comparison with previous years is limited. however, for our westerly and southerly stations we can confirm that the monthly mean temperatures in 2010 were mostly higher than those of previous years (fig. 3). record-setting 2010 in southern greenland the longest running geus measurement series on ice started in 2001 on the qassimiut lobe in southern greenland. this locality was incorporated in the promice network by establishing station qas_l in 2007. the nearly 10 years 1 2 3 4 5 7 8 5 9 7 6 10 11 fig. 2. promice automatic weather station upe_l photographed on 17 august 2009. 1: radiometer. 2: inclinometer. 3: satellite antenna. 4: anemometer. 5: sonic height rangers. 6: thermometer and hygrometer. 7: pressure transducer. 8: solar panel. 9: data logger, barometer and gps. 10: battery box with 4 × 28 ah batteries. 11: 8-level thermistor string. station latitude longitude elevation start date name (°n) (°w) (m) kpc_l* 79°55´ 24°05´ kpc_u 79°50´ 25°10´ sco_l 72°14´ 26°49´ sco_u 72°24´ 27°15´ tas_l 65°38´ 38°54´ tas_u 65°42´ 38°52´ qas_l 61°02´ 46°51´ qas_u 61°11´ 46°49´ nuk_l 64°29´ 49°32´ nuk_u 64°30´ 49°16´ upe_l 72°54´ 54°18´ upe_u 72°53´ 53°32´ thu_l 76°24´ 68°16´ thu_u 76°25´ 68°09´ table 1. promice automatic weather station metadata (status 2010) *l: lower station, u: upper station. 17 july 2008 17 july 2008 21 july 2008 21 july 2008 23 august 2007 15 august 2007 24 august 2007 7 august 2008 20 august 2007 20 august 2007 17 august 2009 17 august 2009 9 august 2010 9 august 2010 380 870 470 1000 270 580 310 890 560 1140 230 980 570 770 75 of data from this locality provide an opportunity to put 2010 into a longer temporal perspective and assess how extraordinary 2010 was at this place. figure 4a shows all available monthly mean temperatures measured at the qas_l site. the qaqortoq temperature record from 56 km southeast of qas_l are included to help interpret the months in 2010 with data gaps due to logger failure (values are reduced by 3°c to facilitate comparison). the climate at qas_l is mild in terms of temperature compared to most other regions on the ice sheet (fig. 3). the lowest winter values do not drop much below –10°c; typical monthly mean winter temperatures during the past decade were in the –10 to –3°c range. summer (june–august) temperatures are predictable in that their mean value is within 2°c of all other years. the presence of ice limits the near-surface air temperature to about 5°c even during warm summers such as 2003. we therefore assume that the august 2010 temperature did not greatly exceed this value, even though the qaqortoq value for this month is the highest ever recorded value (10.6°c). in our records (supported by qaqortoq data), we see that 2010 had the highest on-ice mean temperatures for all months of the year compared to earlier values, with the exception of april (warmer in 2008), july (warmer in 2003, 2005 and 2009) and october (warmer in 2003). this is in full agreement with the values from qaqortoq, which show above-decade average temperatures for all months except july. the most extreme values (in order of excess) were november, may, august, december and september, which exceeded the two standard deviation ranges for the 2000–2009 averages. qaqortoq was on average an astonishing 2.0°c warmer in 2010 than in the second (2003) and third (2005) warmest years on record, 4.5 standard deviations above the 2000–2009 average (which is the warmest decade on record). fig. 3. monthly mean temperatures measured at seven weather station pairs. the upper stations of each pair record the lower temperatures, and vice versa. for locations see fig. 1. 2001 2002 2003 2004 2005 2006 2007 2008 2009 2010 temperatures at qaqortoq (reduced with 3°c). –10 5 –5 0 –10 0 –2 –4 –6 –8 month month 5 6 7 8 9 100 2 4 6 8 10 12 t em pe ra tu re ( ºc ) a b fig. 4. a: monthly mean temperatures and b: cumulative net ablation at lower qassimiut station (qas_l). the ablation measurements are by pressure transducer; supported and validated by sonic ranger where available. the black dot in b shows the total ablation by november 2010. variability in the pressure transducer output is caused by atmospheric pressure. –30 –20 –10 0 kpc sco tas qas nuk upe thu 2008 2009 2010 2011 t e m p e ra tu re ( ºc ) year 7676 even though qas_l summer temperatures are dampened by the ice surface and on average do not exceed 5°c, this does not imply that melt rates are similar between years. the energy consumed by the ice surface to cool the near-surface atmosphere (sensible heat flux) will be larger during warmer periods, as will the down-welling longwave radiation, thus enhancing melting. however, there is only a relatively small amount of year-to-year variability in melt rates (given the slopes of the ablation curves from the pressure transducer in fig. 4b) since solar radiation is the main contributor to melt energy (van as et al. 2009). more important to net abla tion is the length of the ice-melt season, which largely depends on the duration of the period with positive temperatures and the amount of snow accumulation in the preceding winter. for instance, even though the melt rate in 2003 was above average due to high temperatures, the total ablation was near average because of the time it took to melt the relatively large amount of snow that had accumulated during the preceding winter. the year 2005 had lower summer temperatures, but a larger ablation total as there was very little snow accumulation the previous winter. the net ice ablation observations for the period 2001–2009 range from 5 m to 6.5 m of ice per year, which are the largest ablation totals measured anywhere on the greenland ice sheet. for 2010 the extreme months of august and september are lacking from our data series, but spring values show that hardly any snow had accumulated in winter, and that the melt of the bare ice surface began in early may, 1–2 months earlier than in previous years. melt rates were high in late summer and autumn, setting a new ablation record with a measured end-of-year total of about 9 m of ice (fig. 4b, black dot). similar record setting ablation is expected to have taken place in all of southern greenland, as well as along the western margin of the greenland ice sheet (tedesco et al. 2011). conclusions promice has been successful in acquiring near-surface meteorological data over the greenland ice sheet since 2007. temperature measurements display distinct differences between the locations due to solar influences, elevation and regional climate. the promice temperature record confirms that 2010 was an exceptionally warm year in the southern and western regions of greenland, although a longer time series is needed to quantify the 2010 anomaly over the ice sheet. a record-setting net ablation of 9 m of ice in greenland was measured on the southernmost part of the ice sheet in 2010. the enhanced down-welling longwave radiation and sensible heat flux due to the high atmospheric temperatures are not the main reason for the large ablation; low snow accumulation in the previous winter and a long melt season are. acknowledgements the programme for monitoring of the greenland ice sheet (promice) is funded by the danish ministry of climate and energy, and is conducted in collaboration with the national space institute (dtu space) and asiaq (greenland survey). the greenland climate research centre (gcrc) co-finances the nuk stations through the freshlink and imglaco projects. references ahlstrøm, a.p. & promice project team 2008: a new programme for monitoring the mass loss of the greenland ice sheet. geological survey of denmark and greenland bulletin 15, 61–64. box, j.e. 2002: survey of greenland instrumental temperature records: 1873–2001. international journal of climatology 22, 1829–1847. ipcc 2007: intergovernmental panel on climate change (ipcc) fourth assessment report (ar4), climate change 2007. 4 volumes. cambridge: cambridge university press. khan, s.a., wahr, j., bevis, m., velicogna, i. & kendrick, e. 2010: spread of ice mass loss into northwest greenland observed by grace and gps. geophysical research letters 37, l06501. doi: 10.1029/2010gl042460. rignot, e., velicogna, i., van den broeke, m.r., monaghan, a. & lenaerts, j. 2011: acceleration of the contribution of the greenland and antarctic ice sheets to sea level rise. geophysical research letters 38, l05503. doi:10.1029/2011gl046583. tedesco, m., fettweis, x., van den broeke, m.r., van de wal, r.s.w., smeets, c.j.p.p., van de berg, w.j., serreze, m.c. & box, j.e. 2011: the role of albedo and accumulation in the 2010 melting record in greenland. environmental research letters 6, 014005. doi: 10.1088/17489326/6/1/014005. van as, d., bøggild, c.e., nielsen, s., ahlstrøm, a.p., fausto, r.s., podlech, s. & andersen, m.l. 2009: climatology and ablation at the south greenland ice sheet margin from automatic weather station observations. the cryosphere discussions 3, 117–158. van den broeke, m., bamber, j., ettema, j., rignot, e., schrama, e., van de berg, w.j., van meijgaard, e., velicogna, i. & wouters, b. 2009: partitioning recent greenland mass loss. science 326, 984–986. velicogna, i. 2009: increasing rates of ice mass loss from the greenland and antarctic ice sheets revealed by grace. geophysical research letters 36, l19503. doi: 10.1029/2009gl040222. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: dva@geus.dk geological survey of denmark and greenland bulletin 23, 2011, 77–80 77 in the following we describe the project geoscience documents and data for exploration in greenland, in short dodex. a central part of dodex is an interactive web application (http://www.geus.dk/dodex/) that provides easy access to all non-confidential company geoscience reports received by the authorities in greenland and denmark in accordance with the mineral resources act of greenland (1 january 2010) and associated regulations. from the web application it is possible to search in the dodex report database using alphanumeric and geographic search criteria and to access report metadata. it is also possible to download the actual report as a pdf file. in addition to the open dodex web application, the project also includes the development of a closed web application where authorised users can access confidential reports. the dodex project was carried out at the geological survey of denmark and greenland (geus) in cooperation with the bureau of minerals and petroleum (bmp) under the government of greenland as part of the promotion of the mineral resources of greenland. data handling and database the mineral resources act of greenland stipulates that companies holding licences for exploration or exploitation must submit reports on and data from their activities in greenland to bmp. these reports are forwarded to geus where they are scanned and entered into the database. most new reports are confidential for a specific period of time. in dodex – geoscience documents and data for exploration in greenland peter riisager, mikael pedersen, mette svane jørgensen, frands schjøth and leif thorning fig. 1. the dodex web application allows users to search the database using alphanumeric or geographical search criteria. here the user searches for reports on gold in pdf files. the research result will include all reports with geo-references located in the marked red rectangle. © geus, 2011. geological survey of denmark and greenland bulletin 23, 77–80. open access: www.geus.dk/publications/bull 7878 the case of exploration licences, after a five-year confidentiality period, or two years after the relinquishment of an exploration licence, the geoscience reports and submitted data will be considered to be in the public domain and become available to all. besides entering new reports as they are received at geus, we also continuously update the database by scanning and geo-referencing existing reports. the geo-references are polygons that represent areas that are treated in the report. many of the reports were loaded into the database from older databases, and for most of those the area of interest was automatically defined as the area covered by the licence. many of these geo-references have later been edited to represent more accurately the locality or localities treated in the report, and this process of narrowing the areas of interest will continue. new reports entered into the database will be geo-referenced at the time of registration as accurately as practically possible. the reports and report metadata are stored in a relational database housed and maintained by geus (tulstrup 2004). the dodex database model is designed to hold quite a long list of report metadata including obvious parameters such as authors, company, year of publication and commodities described, but also more complex parameters such as geographic reference, confidentiality, and status with respect to quality assurance. at the time of writing the dodex database contains 2151 reports, of which 1197 are non-confidential. of the 1197 released reports 784 have an associated pdf file, and 731 are geo-referenced. user and quality control user control is implemented by database roles that define five different user types, listed with increasing privileges: 1. the public users are only allowed to access non-confidential reports that have been quality assured, and are released. public users correspond to users of the publicly available dodex search web application (figs 1, 2) described in the next section. 2. the trusted users are given a username and password that allow them to log into a restricted web applicafig. 2. the report geo-reference is marked by a blue polygon, while other metadata are listed in the information window. 79 tion (fig. 3), where they can view confidential and unreleased reports. trusted users can only view reports and are not allowed to change the database content. 3. the compilers are also allowed to login to a restricted web application where they can upload and edit reports, however, they are not allowed to change the confidentiality status of reports or release the report for public users. 4. the quality controllers have all the compiler’s privileges and in addition the quality controllers can change report confidentiality status and release reports for public users. the main function of the quality controllers is to review the database work performed by the compilers and to control the quality before reports are released. 5. the administrators develop the system, and can create, read, update and delete all dodex reports and metadata. overall, the setup for dodex user management and quality assurance has proved successful, and hence has been copied in several later geus database projects. the dodex public search web application the public search web application that is available at http://www.geus.dk/dodex/ constitutes the central part of the dodex project. it is designed to be interactive and easy to use, providing easy access for interested parties around the world to all non-confidential company geoscience reports on greenland. from the web application it is possible via alphanumeric and geographic search criteria to search the dodex report database (fig. 1). greenland covers a large geographical area, and an important aspect of the project is that reports are geo-referenced to allow searches in the database for reports relevant for a particular part of the island. the geo-reference polygons of the reports are used for searches on a map of greenland (fig. 1). to minimise transmission time, a relatively simple geological map is used on screen in utm zone 24, with search polygons calculated in decimal degrees, which are the units for the coordinates of report polygons. the search will extract from the database all reports with fig. 3. the dodex administration is handled in a separate web application, where authorised users, depending on their type can view, update, edit, and delete reports. a screen view of the application for a quality controller to check the status of a report is shown. 8080 an area of interest overlapping the search polygon. alphanumeric search criteria can be combined with a geographic search (fig. 1). for the individual reports, public users can obtain a long list of report metadata and download the actual report as a pdf file. it is also possible to add reports to a shopping cart and subsequently get several reports sent by e-mail. each search is logged and stored in the database. since its start in 2008, we have registered a total of 3271 queries from 212 individual users. most often queries are based on geographic location, with north greenland and south-west greenland being the two most popular regions. the two hitherto most sought after commodities are gold and niobium. dodex administration the dodex administration web application is developed for three user types: trusted users, compilers and quality controllers. the administration web application demands login username and password. security is implemented in the web application. it also uses a dedicated one-to-one connection to the database, adding a second layer of security provided by the underlying oracle database authentication and authorisation procedures. this means that when a dodex user logs into the web application the user will also log into a dedicated oracle account. at database level, the authorisation ensures that the users are able to select only data they have the privilege to see, and to create, update and delete. the administration web application is designed to make it easy for the various user types to perform their individual tasks, for example providing the quality controller with an overview of the status of a report (fig. 3). concluding remarks the joint geus-bmp dodex project supports geus’ obligations, described in the 2008–2011 ministerial contract, to make relevant geoscientific data from denmark and greenland accessible to the general public and private companies. the project is in line with geus’ ambitions to embrace the ongoing development in information technology, and, in particular, the advance in web technology to develop more effective tools to make the large quantities of scientific data we hold in our databases available to a broader public (e.g. riisager et al. 2010; hansen & pjetursson 2011 – this volume). the dodex project represents the first attempt at geus to use new open-source java frameworks such as hibernate, and various java server faces frameworks. these steps to modernise software development at geus have proved successful and are now incorporated in its database and software group. most importantly, the dodex project has resulted in a new database for geoscience reports on greenland, and web tools for geus personnel to update and extend this database, and public users to see and download non-confidential reports. hence, dodex has reduced much of the work at geus with organising reports, and photocopying and mailing reports to interested parties, and at the same time made the reports more readily available to the outside world. references hansen, m. & pjetursson, b. 2011: free, online danish shallow geological data. geological survey of denmark and greenland bulletin 23, 53–56. riisager, p., keulen, n., larsen, u., mclimans, r.k., knudsen, c. & tulstrup, j. 2010: interactive web analysis and presentation of computercontrolled scanning electron microscopy data. geological survey of denmark and greenland bulletin 20, 103–106. tulstrup, j. 2004: environmental data and the internet: openness and digital data management. geological survey of denmark and greenland bulletin 4, 45–48. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: pri@geus.dk geological survey of denmark and greenland bulletin 30, 150 pp. 1 geological survey of denmark and greenland bulletin 30 • 2013 stratigraphic landscape analysis, thermochronology and the episodic development of elevated, passive continental margins paul f. green, karna lidmar-bergström, peter japsen, johan m. bonow and james a. chalmers geological survey of denmark and greenland danish ministry of climate, energy and building 22 geological survey of denmark and greenland bulletin 30 keywords apatite fission-track analysis, base level, continental margins, denudation, geomorphology, great escarpment, peneplain, stratigraphic landscape analysis, thermochronology, uplift. cover illustration gåseland, east greenland (70°n, 28°w), looking north-west. the dominant 1900-m plateau is an erosion surface that truncates palaeogene basalts which cover the undulating, weathered basement in the foreground. after uplift which started around the eocene–oligocene transition, the plateau surface was formed by fluvial incision and slope processes, and ultimately graded to sea level. the landscape was then uplifted again in the miocene and incised by rivers. the present plateau is smoothed by periglacial processes and the river valleys have been widened and deepened by glacial erosion. (bonow et al. 2014, in press). frontispiece: facing page planalto da conquista, atlantic margin of brazil (16°s, 42°w, about 200 km from top to bottom). the two low-relief surfaces at c. 900 m (red) and 300 m (green) above sea level are erosion surfaces graded to the base level of the adjacent ocean. the higher surface is palaeogene and was uplifted to its present elevation during the miocene, after which the lower surface formed by incision below the uplifted, higher surface (japsen et al. 2012b). chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, geological institute, university of aarhus; minik rosing, geological museum, university of copenhagen; finn surlyk, department of geography and geology, university of copenhagen scientific editor of this volume: adam a. garde editorial secretaries: jane holst and esben w. glendal referees: paul a.m. andriessen (nl), sierd cloetingh (nl) and piotr migoń (pl) illustrations: stefan sølberg digital photographic work: benny schark layout and graphic production: kristian rasmussen printers: rosendahls-schultz grafisk a/s, albertslund, denmark manuscript received: 21 june 2013 final version approved: 19 november 2013 printed: 30 december 2013 issn 1604-8156 isbn 978-87-7871-372-8 citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 30, 150 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: s and at www.geus.dk/publications/bull © de nationale geologiske undersøgelser for danmark og grønland (geus), 2013 for the full text of the geus copyright clause, please refer to www.geus.dk/publications/bull http://www.geus.dk/publications/bull http://www.geus.dk/publications/bull 3 44 contents abstract . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7 1. introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9 2. geomorphological concepts related to the evolution of elevated, passive continental margins (epcms) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 11 2.1 peneplains as key data for understanding denudation, subsidence and uplift of epcms 11 2.2 historical review of the identification and use of peneplains in landscape analysis . . 14 2.3 criticism of landscape analysis and the idea of steady state in geomorphology . . . . . 16 2.4 different approaches to geomorphological studies of passive margins . . . . . . . . . . 16 3. stratigraphic landscape analysis (sla): a new approach for extracting histories of denudation, subsidence and uplift . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19 3.1 observations and techniques of sla . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19 3.1.1 the sub-cambrian peneplain . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19 3.1.2 the south swedish dome . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19 3.1.3 cross-cutting relationships and their implications . . . . . . . . . . . . . . . . . . . 22 3.1.4 peneplains as unconformities in other areas and their crosscutting relationships 24 3.1.5 sea level, a major base level . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 24 3.1.6 other implications of re-exposed peneplains . . . . . . . . . . . . . . . . . . . . . . . 24 3.1.7 peneplain formation – discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 24 3.2 terminology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 27 3.2.1 landscape, relief and topography . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 27 3.2.2 peneplanation or pediplanation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 27 3.2.3 pediplains (rock-cut plains), pediments, etched surfaces (hilly relief), inselberg plains and etchplains . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 27 3.2.4 classification of peneplains . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 29 3.2.5 high-level benchlands (stepped peneplains), incised valleys and the great escarpment . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 30 3.3 palaeoplains, preservation, destruction and age . . . . . . . . . . . . . . . . . . . . . . . . . 30 3.3.1 effects of glaciation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 32 3.4 combining sla and thermochronology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 33 4. low-temperature thermochronology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 34 4.1 appatite fission-track methods . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 34 4.1.1 historical background . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 34 4.1.2 thermal response of fission tracks in apatite . . . . . . . . . . . . . . . . . . . . . . . 36 4.1.3 variation in annealing kinetics between different apatite species . . . . . . . . . . 37 4.1.4 extracting thermal history information from apatite fission-track data . . . . . 39 4.1.5 monotonic cooling vs. episodic heating and cooling . . . . . . . . . . . . . . . . . . 47 4.1.6 the meaning of a fission-track age . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 48 4.1.7 ‘boomerang plots’ . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 49 4.1.8 uplift rates from apatite fission-track age profiles . . . . . . . . . . . . . . . . . . . . 50 4.1.9 long term residence in the partial annealing zone vs. heating and cooling . . . 50 4.1.10 limitations of apatite fission-track methods . . . . . . . . . . . . . . . . . . . . . . . 51 4.1.11 alternative views of fission-track annealing . . . . . . . . . . . . . . . . . . . . . . . 51 4.1.12 summary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 52 4.2 apatite (u-th)/he dating . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 52 5 4.2.1 historical background . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 52 4.2.2 early success . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 52 4.2.3 evidence of greater complexity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 53 4.2.4 he-closure temperatures and ‘cooling ages’ . . . . . . . . . . . . . . . . . . . . . . . . 55 4.3 converting thermal history information to denudation history . . . . . . . . . . . . . 55 4.3.1 denudation histories . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 55 4.3.2 palaeogeothermal gradients and removed section . . . . . . . . . . . . . . . . . . . 55 4.3.3 elevated heat flow on continental margins . . . . . . . . . . . . . . . . . . . . . . . . 56 4.4 other methods for constraining removed section and denudation histories . . . . . . 57 4.4.1 introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57 4.4.2 vitrinite reflectance . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57 4.4.3 zircon fission-track analysis and zircon (u-th)/he dating . . . . . . . . . . . . . 57 4.4.4 estimating palaeoburial using sonic velocity data . . . . . . . . . . . . . . . . . . . . 58 4.5 constraints from basic geological data and stratigraphic landscape analysis . . . . . . 58 4.6 summary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 59 5. the west greenland margin; a consistent synthesis of geological data, stratigraphic landscape analysis and low-temperature thermochronology . . . . . . . . . . . . . . . . . 60 5.1 an integrated approach . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 60 5.2 geological background: the nuussuaq basin and the offshore record . . . . . . . . . . 60 5.2.1 onshore exposures of the nuussuaq basin . . . . . . . . . . . . . . . . . . . . . . . . 60 5.2.2 the geological record off southern west greenland . . . . . . . . . . . . . . . . . . 67 5.3 results of sla in west greenland . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 68 5.4 use of palaeothermal methods to define an absolute chronology of uplift and exhumation events . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 73 5.5 discussion of thermochronological data from west greenland . . . . . . . . . . . . . . . 77 5.6 integration of palaeothermal data with sla . . . . . . . . . . . . . . . . . . . . . . . . . . . 77 5.7 correlation with the offshore record . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 83 5.7.1 the nature of the oligocene hiatus offshore southern west greenland . . . . . 83 5.8 tectonic evolution of the west greenland epcm . . . . . . . . . . . . . . . . . . . . . . . 86 6. contrasting views on the development of epcms in other areas from landscape studies and thermochronology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 87 6.1 introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 87 6.2 southern africa . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 87 6.2.1 classic landscape studies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 87 6.2.2 low-temperature thermochronology studies . . . . . . . . . . . . . . . . . . . . . . . 88 6.2.3 attempts to reconcile landscape studies and thermochronology in southern africa . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 90 6.2.4 possible alternative explanations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 92 6.2.5 evidence for post-breakup subsidence and burial of the southern africa margin 94 6.2.6 summary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 95 6.3 south-east australia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 95 6.3.1 low-temperature thermochronology studies . . . . . . . . . . . . . . . . . . . . . . . 95 6.3.2 landscape analysis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 96 6.3.3 attempts to reconcile thermochronology and landscape analysis in south-east australia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 97 6.3.4 evidence for episodic burial and uplift on the south-east australian margin. . 99 66 6.4 brazil . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 99 6.4.1 introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 99 6.4.2 studies focused on landscape evidence . . . . . . . . . . . . . . . . . . . . . . . . . . 100 6.4.3 studies focused on thermochronology . . . . . . . . . . . . . . . . . . . . . . . . . . 104 6.4.4 integration of approaches . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 105 6.4.5 continuing controversy . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 109 6.4.6 summary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 110 6.5 other areas . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 111 6.6 numerical modelling of epcm development . . . . . . . . . . . . . . . . . . . . . . . . . 112 6.7 concluding remarks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 112 7. what processes drive the formation of epcms? . . . . . . . . . . . . . . . . . . . . . . . . . . 114 7.1 hypotheses that apply to one or only a few epcms . . . . . . . . . . . . . . . . . . . . . . 114 7.2 permanently uplifted rift margins? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 115 7.3 uplift by underlying hot and/or upwelling mantle . . . . . . . . . . . . . . . . . . . . . . . 117 7.4 uplift due to compression . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 118 7.5 summary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 119 8. key issues concerning the development of epcms . . . . . . . . . . . . . . . . . . . . . . . 121 8.1 steady-state or transient landscapes, ancient landscapes or young landscapes? . . . . 121 8.2 landscape evolution and thermochronology: progressive emergence (continual cooling) vs. episodic burial and exhumation (heating and cooling) . . . 123 8.3 epcms: permanent highs or the result of late uplifts? . . . . . . . . . . . . . . . . . . . . 125 8.4 the meaning of ‘escarpment’ . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 127 8.5 down-wearing, back-wearing and scale . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 127 8.6 are epigene surfaces sometimes re-exposed? . . . . . . . . . . . . . . . . . . . . . . . . . . . 129 8.7 episodic development of epcms . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 129 8.8 observations vs. theory in understanding the evolution of epcms . . . . . . . . . . . . 130 9. summary and conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 132 acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 132 references . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 133 7 abstract green, p.f., lidmar-bergström, k., japsen, p., bonow, j.m. & chalmers, j.a. 2013: stratigraphic landscape analysis, thermochronology and the episodic development of elevated, passive continental margins. geological survey of denmark and greenland bulletin 30, 150 pp. the continental margin of west greenland is similar in many respects to other elevated, passive continental margins (epcms) around the world. these margins are characterised by extensive regions of low relief at elevations of 1–2 kilometres above sea level sloping gently inland, with a much steeper, oceanward decline, often termed a ‘great escarpment’, terminating at a coastal plain. recent studies, based on integration of geological, geomorphological and thermochronological evidence, have shown that the high topography of west greenland was formed by differential uplift and dissection of an oligo-miocene peneplain since the late miocene, many millions of years after continental break-up between greenland and north america. in contrast, many studies of other epcms have proposed a different style of development in which the high plateaux and the steep, oceanward decline are regarded as a direct result of rifting and continental separation. some studies assume that the elevated regions have remained high since break-up, with the high topography continuously renewed by isostasy. others identify the elevated plains as remnants of pre-rift landscapes. key to understanding the development of the west greenland margin is a new approach to the study of landforms, stratigraphic landscape analysis, in which the low-relief, high-elevation plateaux at epcms are interpreted as uplifted peneplains: low-relief surfaces of large extent, cutting across bedrock of different age and resistance, and originally graded to sea level. identification of different generations of peneplain (re-exposed and epigene) from regional mapping, combined with geological constraints and thermochronology, allows definition of the evolution leading to the formation of the modern-day topography. this approach is founded particularly on results from the south swedish dome, which document former sea levels as base levels for the formation of peneplains. these results support the view that peneplains grade towards base level, and that in the absence of other options (e.g. widespread resistant lithologies), the most likely base level is sea level. this is particularly so at continental margins due to their proximity to the adjacent ocean. studies in which epcms are interpreted as related to rifting or break-up commonly favour histories involving continuous denudation of margins following rifting, and interpretation of thermochronology data in terms of monotonic cooling histories. however, in several regions, including southern africa, south-east australia and eastern brazil, geological constraints demonstrate that such scenarios are inappropriate, and an episodic development involving post-breakup subsidence and burial followed later by uplift and denudation is more realistic. such development is also indicated by the presence in sedimentary basins adjacent to many epcms of major erosional unconformities within the post-breakup sedimentary section which correlate with onshore denudation episodes. the nature of the processes responsible is not yet understood, but it seems likely that plate-scale forces are required in order to explain the regional extent of the effects involved. new geodynamic models are required to explain the episodic development of epcms, accommodating post-breakup subsidence and burial as well as subsequent uplift and denudation, long after break-up which created the characteristic, modern-day epcm landscapes. authors’ addresses p.f.g., geotrack international, 37 melville road, brunswick west, victoria 3055, australia. corresponding author’s e-mail: mail@geotrack.com.au k.l.-b., department of physical geography and quaternary geology, stockholm university, se-106 91 stockholm, sweden p.j. & j.a.c., geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350 copenhagen k, denmark j.m.b., geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350 copenhagen k, denmark and södertörn university, se-141 89 huddinge, sweden mailto:mail@geotrack.com.au 88 70˚w fig. 2 fig. 4 fig. 5 fig. 3 elevated, passive continental margins 5 9 elevation (km a.s.l.) 3210-1.5-3-5-7-11 fig. 1. topography of the earth between 80°n and 60°s showing elevated, passive continental margins (epcms). only atlantic-type margins that can easily be connected to the corresponding spreading centre and only elevated margins that reach 2 km a.s.l. in more than a single summit are indicated (cf. bradley 2008). global relief model: http://www.ngdc.noaa.gov/mgg/image/images/g01929-pos-a0001small.pdf. modified from japsen et al. (2012a). http://www.ngdc.noaa.gov/mgg/image/images/g01929-pos-a0001small.pdf 9 1. introduction the topography and relief of the continental margin of west greenland are similar in many respects to those of other rifted continental margins around the earth. margins in locations as diverse as south-east australia, southern africa, brazil, norway, western india, saudi arabia and yemen (fig. 1) are characterised by inland regions of low relief up to 1 or 2 km or more above sea level (a.s.l.), declining gently inland and separated from a narrow coastal plain by a region of much steeper oceanward decline (jessen 1948). the relief of these elevated, passive continental margins (epcms) is characterised by high level plains, commonly in stepped sequences, with deeply incised valley systems (figs 2–5). at many margins, mesozoic–cenozoic rift systems parallel to the coast are present offshore with a transition from continental to oceanic crust farther offshore. post-rift sedimentary sequences offshore commonly dip away from the coast and are truncated below unconformities which correlate with onshore denudation episodes (fig. 6; japsen et al. 2012a and references therein). despite many years of intensive study, the tectonic development of epcms is still only poorly understood, particularly in regard to the timing and magnitude of uplift, amounts and timing of denudation (‘missing section’) and the nature of the underlying processes. since the 1950s and 1960s there has been much dispute about the origin of the high plains as peneplains graded to former sea levels, and the idea has passed out of favour among many geomorphologists (phillips 2002). many studies are based on the assumption that epcm morphology is directly related to the processes of rifting, continental break-up and separation, and that epcms have remained high since prior to break-up, as a result of continuous erosion and isostatic adjustment (e.g. gilchrist & summerfield 1991; gallagher & brown 1997; gallagher et al. 1998; bishop 2007). others (e.g. ollier & pain 1997) have identified the elevated plains as remnants of pre-rift landscapes which have remained largely unaffected by later processes. in contrast, recent studies of the west greenland margin have shown that the present-day mountains, with plateaux up to 2 km or more a.s.l., represent a differentially uplifted, broken and dissected planation surface graded to sea level during the oligo–miocene and uplifted in the late miocene and pliocene (japsen et al. 2005, 2006, 2009; bonow et al. 2006a, b). following paleocene breakup between greenland and north america, sea-floor spreading ceased at the end of the eocene (chalmers & pulvertaft 2001), so the uplift of these plateaux postdates rifting and continental separation by many millions of years, and the present-day topography cannot be related to continental break-up. since the west greenland margin shares many of the characteristics common to ecpms worldwide, these results seriously question the common belief that the elevated plateaux with oceanward escarpments characteristic of epcms are directly related to continental rifting and separation. a cornerstone of these studies in west greenland has been the combination of observations from the onshore geological record with a) an improved methodology for landscape analysis (stratigraphic landscape analysis, sla, see chapter 3), which defines a relative chronology of denudation, subsidence and uplift, and b) thermochronology incorporating analysis of both basement and overlying sedimentary cover rocks, which provides an absolute timescale and magnitude for key denudational and depositional events (chapter 4). the combination of these methods results in a consistent definition of an episodic style of evolution, involving multiple episodes of both positive and negative vertical movements. integration of the resulting history with additional information from geological constraints offshore (chalmers 2000; japsen et al. 2005, 2006; bonow et al. 2007b) has provided a coherent regional framework for deciphering the post-rift tectonic development of the west greenland epcm, as explained in chapter 5. in contrast to this consistent fusion of stratigraphic landscape analysis and thermochronology in west greenland, many studies of epcms in other parts of the world have revealed a conflict between conclusions derived from conventional landscape studies and thermochronology (e.g. ollier & pain 1997; peulvast et al. 2008). any viable model for the development of epcms should be consistent with all available constraints. chapter 6 describes studies of a number of other epcms (including southern africa, south-east australia and brazil) where previous studies have led to conflict between landscape studies and thermochronology, and proceeds to illustrate how the different approaches can be integrated to define an episodic style of development similar to that established for the west greenland margin in chapter 5. chapter 7 deals with the nature of the underlying pro1010 cesses, and possible mechanisms which might explain the episodic development of epcms. several key issues regarding the development of epcms are highlighted and discussed in chapter 8. this study will particularly focus on two major problems in previous studies of epcms. the first is the notion of permanently uplifted margins and the paradigm of steady state (uplift due to isostasy keeps pace with erosion; see section 2.3) which for a long time prevented the use of peneplains as uplift markers in geomorphology. the second is the common adoption in thermochronological studies of monotonic cooling histories. basic constraints provided by geological observations which impose critical restrictions on viable solutions have commonly been overlooked in both geomorphological and thermochronological studies. the evidence presented here leads to the conclusion that the high present-day elevation of many passive continental margins is not directly related to processes of continental rifting and break-up, but to processes taking place subsequently. the common assumption that epcms are permanently elevated cannot be sustained, and a reassessment of the development of epcms is required. the constraints that can be derived from thermochronology and landscape analysis, integrated with the geological record, can provide the basis for such reassessment. a detailed understanding of the development of epcms is important not only in terms of understanding the properties of the earth’s lithosphere and mantle but also has major implications for hydrocarbon exploration, e.g. in terms of sediment supply to offshore basins. lower jurassic basalt d r a k e n s b e r g 0 500 1500 2000 elevation (m a.s.l.) 2500 30001000 n 50 km 67006600 6800 6900 7000 1000 900 800 700 600 500 fig. 2. south africa, south-east coast, with the drakensberg escarpment in lower jurassic basalt. below it are several planation surfaces separated by major steps (cf. king 1972) including the african surface at about 1800 m a.s.l., a major plain at about 1000 m a.s.l., and a coastal plain. two valley generations are incised below the 1000 m planation surface. the general picture of the relief at the elevated, passive margin of south africa is of high-level plains and deeply incised valleys. the map covers most of the kwazulu-natal province. 3d-terrain model with vertical shading. utm coordinates (km) zone 36s. elevation data from jarvis et al. (2008). 11 2. geomorphological concepts related to the evolution of elevated, passive continental margins (epcms) 2.1 peneplains as key data for understanding denudation, subsidence and uplift of epcms the relief of epcms (figs 2–5) commonly takes the form of high-level plains, separated by more or less pronounced steps. a gentle overall slope characterises their inland continuation, while the oceanward slope is steeper and can take a variety of forms, including distinctly separated steps (fig. 2), inclined plains (figs 3, 4) or tilted fault blocks (fig. 5). in all cases, deeply incised valleys are present. what these margins specifically have in common are the high-level plains and incised valleys, which according to traditional landscape analysis suggest late uplift events, as the plains have been interpreted as having formed close to sea level (davis 1899; ahlmann 1919, 1941; reusch 1901; johnson 1931; king 1951, 1962, 1967, 1972, 1976; peulvast 1985;, 1987; partridge & maud 1987). the margins can show tilted surfaces (e.g. australia, fig. 3), where the incised valleys have coalesced to form a great escarpment (ollier 1982). the notion of a great escarpment has been used to characterise a variety of margins (e.g. ollier 1985; ollier & marker 1985; partridge & maud 1987; lidmar-bergström et al. 2000; gunnell et al. 2003) although the so-called great escarpments greatly differ in shape in different settings (see above). the inclined plains characterising some margins can be interpreted as re-exposed unconformity surfaces (figs 3, 4; section 3.3) and suggest the former presence of more extensive covers. the high plains and the re-exposed unconformities are important data for revealing histories of denudation, subsidence and uplift (e.g. lidmar-bergström 1993; bonow 2005; bonow et al. 2006a; lidmar-bergström et al. 2013). although glaciation produces specific features such as deep fjords and erases summit plains in locations exposed to cirque and valley glaciation, which form an alpine relief (figs 4, 5; lidmar-bergström et al. 2000; mitchell & montgomery 2006; bonow et al. 2006b; etzelmüller et al. 2007), the high plains and incised valleys are common features at all epcms, whether they are glaciated or not. low-relief landscapes hundreds of kilometres in extent at both high and low elevations are common features on the earth’s continents (e.g. king 1962, 1967, 1976; bird 1959; ollier 1981; hall 1991; de brum ferreira 1991; lidmar-bergström 1996; gunnell 1998; johansson 1999; godard et al. 2001b; lageat & gunnell 2001; casa-sainz & cortés-gracia 2002; demoulin 2003; bonow et al. 2003, 2006a). they were labelled peneplains and used at the beginning of the 20th century to constrain base-level changes and uplift (e.g. davis 1899; reusch 1901; ahlmann 1919; johnson 1931). the following principles were applied: 1) large areas of low relief at high levels forming subhorizontal or tilted planes across different rock types are produced by erosion to a common base level (the sea, in near-shore positions) and have subsequently been uplifted to their present elevations; 2) scattered summit surfaces at higher elevations bevelling different rock types may be remnants of older peneplains; 3) valley incision to common lower levels shows younger generations of erosion graded to new base levels. if the peneplains cannot be related to any covers and have thus been exposed since formation, they are labelled epigene (twidale 1985). since the definition in the late 19th century of a peneplain as the result of fluvial incision and ultimate grading to base level, knowledge about deep weathering processes and resulting bedrock shapes arose in the 1950s (section 3.2). it is now also known that some peneplains have been preserved below cover rocks of different age (section 3.1.2). therefore a new classification of peneplains is presented in this paper (section 3.2.4.) which embraces low relief surfaces with different shapes and saprolites characteristics. neither the result of deep weathering nor the importance of former covers have been included in previous discussion of peneplains. in the prevailing paradigm of steady state and dynamic equilibrium (inkpen 2005) the existence of peneplains is regarded as elusive (phillips 2002). on the contrary, here the high plateaux are interpreted as peneplains (see section 3.1), which we define as low-relief erosion surfaces graded to distinct former base levels, irrespective of their detailed development. the elevated peneplains can be labelled transient landscapes (bishop 2007), as they are under destruction by valley incision. 1212 grafton ip ip ip high-level plains a b 0 500 1000 1500 2000 2500 3000 25 km high escarpment low escarpment mesozoic sediment townip inclined plains miocene basalt basement limit of inclined plains elevation (m a.s.l.) 350 450 550 6350 6450 6550 6650 6750 n fig. 3. eastern australia. a: 3d-terrain model with vertical shading. b: line drawing of the map in a. there is a major plateau at 800–1200 m a.s.l. and incised valleys coalescing to the great escarpment and a coastal plain. the escarpment is lower where it cuts into inclined plains (ip) that are tilted towards the coast. south-west of grafton, such an inclined plain forms the surface of a triangular basement block, which seems to disappear below jurassic sedimentary strata (geology of australia 1976). this tilted and thus probably sub-jurassic surface, is cut off by the high plain, here capped by miocene basalts (johnson 2004). scale bar applies only to the foreground of the 3d model. utm coordinates (km) zone 56s. elevation data from jarvis et al. (2008). facing page: fig. 4. western norway (glaciated). a: 3d-terrain model with vertical shading. b: line drawing of the map in a. there are extensive, high-level plains and valleys incised below about 800 m a.s.l. and a coastal plain, the strandflat. as in parts of eastern australia (fig. 3) a tilted, re-exposed, and probably sub-mesozoic plain can be traced in the summit relief from the high areas to the coast between the major incised valleys (lidmarbergström et al. 2000). areas with major glacial reshaping occur both in high (alpine relief) and low (g) positions. the raundalen valley (r) has escaped major glacial reshaping, while the sognefjord (s) and hardangerfjord (h) have been deepened and widened by major outlet glaciers. as a consequence their inner parts have been fluvially rejuvenated during the interglacials. the møre-trøndelag fault complex (mtfc) is a tectonic zone along which movements occurred in the mesozoic and maybe also later (redfield et al. 2005). jurassic outlier in the bergen area (fossen et al. 1997). scale bar applies only to the foreground of the 3d model. utm zone 32n. elevation data from jarvis et al. (2008). 13 c o a s t a l p l a i n gg gg gg h i g h l e v e l p l a i n sm tfc m tfc gg gg gg gg gg s r h gg alpine relief s r h s r h ice cap inclined re-exposed surface g a b a l p i n e r e l i e f 0 500 1000 1500 2000 2500 3000 basement glacial reshaping in low positions river jurassic outlier limit of glacial reshaping elevation (m a.s.l.) 6950 6850 6750 6650 6550 250 350 450 550 n 25 km 1414 the elevation of epigene peneplains graded to former sea levels can be used as a measure of rock uplift since their formation. this notion on absolute uplift should not be confused with ‘surface uplift’ defined by england & molnar (1990) as the change of mean height of an area. 2.2 historical review of the identification and use of peneplains in landscape analysis the fundamental importance of a base level for erosion by the fluvial system was recognised at the end of the 19th century and formulated by davis (1899) in his classic work on ‘the geographical cycle’. from observations of different landscapes the following model was constructed. a cycle started with uplift of a flat landscape, which caused incision of valleys and the formation of a youthful landscape. further development was thought to occur by slope decline to form a landscape with hills and valleys, a mature landscape. the end result was a new peneplain with a few residual hills, called monadnocks. penck (1924) described peneplains in stepped sequences and many authors (e.g. reusch 1901; ahlmann 1919, 1941; johnson 1931) understood them as markers of former base levels, realising that they could be used to determine uplift events. baulig (1928) noted a re-exposed and tilted, mainly sub-jurassic peneplain along the flanks of the massif central in france and a major horizontal summit surface which could be dated to the oligocene by terrestrial deposits. he interpreted this surface as having formed after an uplift event, as it cuts off the inclined sub-jurassic peneplain. in addition, incised younger valley generations were identified, confirming further base-level changes. at that time the continents were considered basically stable but it was realised that during certain time intervals the sea had flooded large parts of the continents. for the northern hemisphere the late cretaceous was such a period. surfaces and valleys in stepped sequences were identified in several european massifs and thought to reflect a changing (lowering) sea level from the early cenozoic and onwards (baulig 1935). this ‘eustatic’ theory (not to be confused with the modern concept used in sequence stratigraphy) had an immense influence on geomorphological thinking in central europe and particularly in great britain. a major theme in geomorphological studies became the identification of as many steps as possible with the aid of topographic profiles. international commissions tried to correlate peneplains across the atlantic (lefèvre 1960). the efforts were not successful and the old eustatic theory fell out of favour. as a result, the concept of the peneplain in landscape analysis became widely criticised and baulig’s identification of re-exposed peneplains and their significance for extracting histories of uplift and subsidence was not appreciated at that time. although continental drift was not generally accepted in the early part of the 20th century, it was clear to some geologists working in southern africa (e.g. du toit 1937). king (1951, 1956b, 1962, 1967) recognised major low-relief erosion surfaces, which he described as pediplains (see section 3.4) formed by valley incision to former base levels and parallel scarp retreat, as providing important constraints on intermittent uplift after continental break-up and formation of a new coastline. king’s model with scarp retreat contrasts to davis’ model with slope decline, but in both models the base level for the erosion surfaces is a marker of uplift. king’s ideas on development of epcms have had considerable influence and are therefore reviewed here. king (1956b, 1972, 1983) concluded from analysis of landforms and geology that the landscape of natal, southern africa (fig. 2), had developed through a series of uplifts by warping of the continental margin since the jurassic, and he dated the pediplains by correlation to offshore or coastal sediments. he recognised small remnants of the pre-break-up gondwana surface on some high summits above the drakensberg escarpment. the major summit plain below this escarpment, the african surface, he regarded as palaeogene with a major erosional phase already in the late cretaceous, but finally shaped in the early miocene. he regarded two major valley generations as evidence of neogene uplift events. king’s (1983) major message was that the present high position of the rims of southern africa is due to neogene uplifts after formation of the african surface close to sea level: “they have nothing to do with the mesozoic break-up and subsequent continental drift; they refer primarily to cenozoic tectonic movements operating solely in the vertical sense, both up and down” (king 1983, pp. 203– 204). king’s view of the importance of cenozoic tectonics seems to have been forgotten in more recent times, although his work is often acknowledged. in contrast, ollier (1982, 1985) described the highlevel plains inland of an epcm as a single ancient ‘palaeoplain’, with its oceanward termination and decline towards the coastal plain defining a ‘great escarpment’ (even though the oceanward slope can look very different in different places, cf. section 2.1). he regarded this pal15 ice cap inclined re-exposed surface town g g g g 0 1 0 1 2 7400745075007550 n s sl sl gg gg g g g g g lps ups ups b el ev at io n (k m a .s. l.) el ev at io n (k m a .s. l.) a a a b c b b' a' a'b' 400 500450 w e high-level plains high-level plains alpine relief high-level plains upper planation surface (ups) lower planation surface (lps) inclined re-exposed etch surface (es) upses kangerlussuaqkangerlussuaq l l l l l l l l l l l l l l l 0 200 400 600 800 1000 1200 1400 1600 1800 2000 2200 2400 elevation (m a.s.l.) 7650 7600 7550 7500 7450 7400 7350 7300 400 450 500 basement fault limit of glacial reshaping glacial reshaping in low positions 25 km n fig. 5. west greenland (glaciated). a: 3d-terrain model with vertical shading. b: topographical profiles with maximum and minimum elevations within 20 km wide zones. c: line drawing of the map in a. there are well-preserved high-level plains inland, deeply incised valleys and a coastal plain (strandflat). the high plain (upper planation surface, ups) is tilted and offset along faults, and a partially developed, lower plain (lower planation surface, lps) is developed in two areas of major uplift. an inclined, mesozoic etch surface (es) is re-exposed at the coast in the north-west, just south of the cretaceous–paleocene nuussuaq basin (see fig. 36). areas with major glacial reshaping (g) occur both in high (alpine relief) and low positions. sl: sisimiut line offsetting the ups (see fig. 34). scale bar applies only to the foreground of the 3d model. utm coordinates (km) zone 22n. elevation data from jarvis et al. (2008). see chapter 5 for uplift history. 1616 aeoplain as an originally high plain of mesozoic age and discussed how downwarp of this palaeoplain after breakup caused valleys to incise and form a ‘great escarpment’ along the passive margins of eastern australia and southern africa. ollier & marker (1985) pointed to the general form of high passive margins (what we now call epcms) and discussed their development as either uplifted mesozoic rift shoulders or downwarped mesozoic surfaces (ollier & pain 1997). ollier and coworkers did not discuss margins in the frame of neogene uplifts or upwarps, which had been suggested earlier, by e.g. reusch (1901) for scandinavia, craft (1933) for southeast australia, ahlmann (1941) for east greenland and king (1956b) for southern africa. the conceptual models of marginal development (downwarp or rift shoulder uplift) by ollier (1985) are of particular importance as they have been used in discussing the thermal and denudation histories of continental margins (as discussed in chapter 6). 2.3 criticism of landscape analysis and the idea of steady state in geomorphology since the 1950s, when a dynamic basis for geomorphology was advocated by strahler (1952), a controversy has existed within geomorphological circles regarding whether or not landscape generations (stepped peneplains and valley generations) graded to distinct former base levels exist. researchers such as chorley (1965) and summerfield (2000) denied their existence, whereas other geomorphologists of the new dynamic school such as schumm (1975) and ahnert (1994) never questioned the old principles of landscape generations. chorley’s (1965) main criticism, however, was aimed at clarifying for geographical geomorphologists that denudation chronology is not a relevant issue “for the geographer with his human theme”, although he admitted that landform analysis in combination with stratigraphy could give scientific results (e.g. wooldridge & linton 1955). yet, his criticism of landform analysis resulted in many geomorphologists being reluctant to use peneplains and valley generations as a scientific data set to reveal uplift events. criticism of landscape analysis has continued within physical geography through the concept of the ‘steady state’ (inkpen 2005), in which the same amount of energy is thought to enter and leave the system without causing any major changes in topography as it is assumed that uplift due to isostasy keeps pace with erosion. this view does not acknowledge the presence of inherited landforms such as high-level peneplains graded to old base levels (cf. hack 1960), and the information that the landscape elements contribute for deciphering the sequence of events is ignored. old textbooks based on davisian geomorphology have been abandoned. literature on peneplains and incised valleys and what they tell about past events came to be regarded as unscientific within geomorphology. studies on erosion rates, based on thermochronology of the bedrock and cosmogenic isotopes produced in situ, have replaced landform analysis as the major theme in studies of long-term landform evolution and uplift of continental margins (e.g. brown et al. 2000; cockburn et al. 2000; gleadow & brown 2000; summerfield 2000; persano et al. 2002; van der beek et al. 2002). relief is different in different settings and it is evident that there is information in the landscape itself that is of importance for solving tectonic problems (e.g. gellert 1990; burbank et al. 1999; sugai & ohmori 1999; calvet & gunnell 2008; gunnell et al. 2009; westaway 2009). as bishop (2007, p. 330) states, post-orogenic, non-steady-state landscapes “deserve much more attention than they have received of late”. such landscapes are of interest not only because of the questions they pose about their formation and preservation but also because of the information they can provide on tectonic events through the presence of e.g. epigene and re-exposed peneplains and their cross-cutting relationships. 2.4 different approaches to geomorphological studies of passive margins the long-term geomorphological development of passive margins has been studied using a variety of different approaches. in the most common present-day approach (e.g. bishop & goldrick 2000; brown et al. 2000; summerfield 2000; persano et al. 2005), the existence of elevated preserved peneplains graded to ancient common base levels is denied. the topography of epcms, described as consisting of a high-level plain with an oceanward great escarpment, approximately at the main divide, is interpreted as a steady state phenomenon. in this approach the existence of a high plain, originally formed prior to break-up (ollier 1982), is accepted, but the question of how this plain was formed is not dis17 0 1 tw ow ay ti m e (s ec ) 2 0 1 tw ow ay ti m e (s ec ) 2 0 1 tw ow ay ti m e (s ec ) tw ow ay ti m e (m se c) 2 sw 20 km 50 km 20 km 10 km jurassi c ne w e w e w e nw d ep th (k m ) eocene–holocene ku–paleocene albian–cenomanian se0 1 2 3 4 5 10 km coniacian quaternary neogene miocene quaternary paleocene–eocene late santonian miocene quaternary mid-albian to turonian 100 200 300 0 100 200 300 400 500 600 d ep th (m ) 10 km a. offshore west greenland c. offshore se brazil b. offshore north-east greenland d. offshore sw africa e. offshore south norway f. offshore se australia kl ku palaeogene pre-jurassic? plio-pleistocene? paleocene– miocene? kl plio-pleistocene cretaceous upe– jur upe– jur de– mpe palaeogene neogene w e sediments deposited after phases of uplift and erosion post-rift, thermal subsidence sediment syn-rift rock basement volcanics post-rift erosional truncation fig. 6. post-rift truncation of sedimentary sequences offshore epcms indicating post-rift subsidence and uplift of these margins. a: offshore west greenland (70°30´n); oligocene and late neogene truncations along low-angle and high-angle unconformities, respectively (see seismic section in fig. 42a; redrawn from chalmers 2000; japsen et al. 2006, 2010 & unpublished afta data). b: offshore north-east greenland (78°n); early and late neogene truncations (rifting continued into the palaeogene in the central part of the profile; redrawn from hamann et al. 2005). c: offshore south-east brazil (26°30´s); mid-eocene truncation (redrawn from cobbold et al. 2001). d: offshore south-west africa (30°s); post-cretaceous truncation (redrawn from stevenson &mcmillan 2004). post-rift exhumation along south-west africa is documented by walford & white (2005). e: offshore south norway (59°n); early and late neogene truncations along low-angle and high-angle unconformities, respectively (the jurassic sequence may include a thin syn-rift section). the early neogene exhumation is documented by japsen et al. (2010). f: offshore south-east australia (34°s); truncations are above pre-rift rocks whereas syn-rift sediments are missing (base-cenozoic and pliocene dating according to davies 1975). note that low-angle unconformities may be especially difficult to identify as erosional truncations, and consequently that not all truncations may have been identified on these profiles. de–mpe: devonian – middle permian. kl: lower cretaceous. ku: upper cretaceous. upe–jur: upper permian–jurassic. modified after japsen et al. (2012a). 1818 cussed. the plateau inland of the margin is regarded as having continuously been a plain which has remained elevated since prior to rifting, and no consideration is given to the possibility that the margin may have been buried below a sedimentary cover before development of the present-day relief. the elevated topography is a basic assumption about initial conditions at the time of rifting, and is used to define a framework for calculations of denudation rates in different landscape settings (the high plain, the great escarpment, the coastal plain) using low temperature thermochronology. a different approach is used by many french researchers. they use the palaeolandforms as a source of data from which conclusions on landscape development can be drawn (e.g. peulvast et al. 1996) and regard the landforms per se as providing information about past processes (e.g. godard et al. 2001a). the inversion problem connected with this approach was discussed in detail by gunnell (1998). this field-based approach acknowledges structural and lithological control (peulvast & vanney 2001) and the existence of peneplains. when integrating with thermochronology data, approaches differ somewhat between researchers. while peulvast et al. (2008) consider large (km-scale) thicknesses of former sedimentary cover as geologically implausible, gunnell (1998) regards such kilometre-scale former sedimentary covers as reasonable. the idea that high-level plains with incised valleys are evidence of late uplift refers to ‘landscapes of youth’ as described by davis (1899). it was adopted by reusch (1901), who referred to the high plateaux of norway in terms of ‘the palaeic surface’, which he thought to include two levels. a similar approach was adopted by ahlmann (1941), who used similar plateaux in east greenland to infer neogene uplift. in later studies, a relationship has been acknowledged between high plains and re-exposed relief on tilted peneplains between the incised valleys (lidmarbergström 1982, 1988). stratigraphic landscape analysis (chapter 3) was developed with the south swedish dome as a key area, as this dome rises only slightly above surrounding cover rocks, making the relationships between relief and cover more evident. stratigraphic observations and cross-cutting relationships were used to construct a relative chronology for surface formation and tectonic events, both uplift and subsidence with formation of temporary covers. this method has then been applied to sweden (lidmar-bergström 1995, 1996), other parts of scandinavia (lidmar-bergström 1999; lidmarbergström et al. 2000; lidmar-bergström & näslund 2002; bonow et al. 2003; lidmar-bergström et al. 2013) and west greenland (bonow 2005; bonow et al. 2006a). it was integrated with thermochronology and other geological observations in further studies of the west greenland margin (japsen et al. 2006, 2009). similar data for constraining late uplift have been presented independently in papers by authors working on uplift in tibet (importance of former base levels: schoenbohm et al. 2004; clark et al. 2006), corsica (importance of both former base levels and covers: kuhlemann et al. 2005) and south america (importance of uplift and tilting for valley incision: schildgen et al. 2007). while thermochronology is used in many approaches, in recent studies of west greenland (japsen et al. 2006, 2009) it was integrated with stratigraphic landscape analysis and other geological observations to place peneplain formation and subsequent uplift within an absolute timeframe. in these studies, these three types of observations were given equal importance in deriving conclusions in regard to past events, paying particular attention to the possibility of re-burial and the influence of former sedimentary cover (subsequently eroded) in preserving old surfaces. in contrast, some other approaches have resulted in interpretations of thermochronology which are at odds with simple geological constraints, as discussed in chapter 6. 19 3. stratigraphic landscape analysis (sla): a new approach for extracting histories of denudation, subsidence and uplift this chapter presents a new approach to the analysis of landscape, stratigraphic landscape analysis (sla; lidmarbergström et al. 2013). we highlight observations on the formation of peneplains close to sea level, suggest how different terms can be used to avoid confusion and describe evidence for preservation and destruction of old peneplains (palaeoplains). subsequent chapters describe how sla can be combined with low-temperature thermochronology (principally afta) to provide quantitative constraints on the timescales and magnitudes of vertical movements involved in these processes. 3.1 observations and techniques of sla this new approach for landscape analysis is based on experience from the baltic shield and the south swedish dome in particular (figs 7, 8; lidmar-bergström 1988; lidmar-bergström et al. 2013). two concepts are important, viz. basement and cover rocks. basement is defined here as denuded metamorphic and granitic rocks of different orogenies, while cover rocks are undeformed sedimentary or volcanic strata resting directly on basement. we further discriminate between orogeny, a rock-forming process, and mountain building, a process creating topography (ollier 1981). 3.1.1 the sub-cambrian peneplain extensive areas of the baltic shield consist of a flat landscape, which has been identified as the sub-cambrian peneplain by the aid of remnants of cover rocks and cambrian fissure fillings (högbom 1910; högbom & ahlström 1924; tanner 1938; mattsson 1962; rudberg 1960, 1970; lidmar-bergström 1996). it extends for 700 km along the east coast of sweden and 450 km across the central swedish lowlands (fig. 7). the sub-cambrian peneplain can be followed as an inclined surface in boreholes and on seismic profiles for more than 400 km and down to over 4000 m below sea level offshore southeastern sweden and in the baltic countries (kornfält & larsson 1987; fredén 1994). the sub-cambrian peneplain continues on land as a re-exposed inclined surface for over c.100 km from about sea level in the east to over 300 m a.s.l. in the west and forms the crest of the south swedish dome (fig. 9a; lidmar-bergström 1988, 1996). the surface is weathered at its contact with its cover rocks (e.g. hadding 1929; lundegårdh et al. 1973). at the contact there is shallow weathering to a maximum of 5 m (references in elvhage & lidmar-bergström 1987) and the whole surface is extremely flat (fig. 9b). thus the precambrian basement, which formed during different orogenies, had been denuded to an almost featureless plain before the cambrian. the peneplain was flooded by shallow seas during the cambrian and ordovician, and sea level fluctuated only a few tens of metres (artyushkov et al. 2000). the relationship between the surface and the cambrian outliers on it shows that the peneplain must have been virtually horizontal and situated close to the sea level before the transgressions. 3.1.2 the south swedish dome the south swedish dome emerges from below cambrian cover rocks in the north (väner basin) and east and mesozoic cover rocks in the south and west and is delimited to the south-west by the sorgenfrei–tornquist zone (figs 7, 8; lidmar-bergström 1988). the cambrian seas flooded all of southern sweden (jaeger 1984), now occupied by the south swedish dome. as both jurassic and cretaceous cover rocks now rest directly on the basement in areas flanking the present dome in the southwest, it is clear that the palaeozoic cover had been eroded here by mesozoic time. cambrian fissure fillings, nonetheless, confirm the former existence of a lower palaeozoic sedimentary cover (martinsson 1968; samuelsson 1975). both hilly forms of the fresh basement, with a relative relief up to 200 m, and thick (up to 60 m) remnants of kaolinitic saprolite characterise the sub-mesozoic peneplains (fig. 10). analysis of the cross-cutting relationships to the sub-cambrian peneplain shows that the hilly submesozoic, mainly sub-cretaceous, relief around the 2020 l l l l fig. 8 ssd v ssp basement: domes high, >1500 m medium high, 1000–1500 m väner basin premontane region and low dome/ridge basement lows: relief differentation sub-cambrian peneplain sub-mesozoic undulating hilly relief plains with residual hills 8°e 60°n 64°n 56°n 16°e 8°e 16°e 24°e cover rocks muddus plains 2097 m 2468 m 377 m n o r t h e r n s c a n d e s s o u t h e r n s c a n d e s muddus plains ö ba y o f b oth nia palaeozoic mesozoic and cenozoic jurassic outlier vätter graben caledonian front 200 km l l l l l l l l l l l l sorgenfrei–tornquist zone c o n t i n e n t a l s l o p e mixed relief b mixed relief a sca ndinavia norway sweden finland fig. 7. topography and relief of basement in scandinavia in relation to the surrounding cover rocks. three domes are identified and their highest parts marked on the map: the northern scandes, the southern scandes and the south swedish dome (ssd). three types of low relief are identified on the flanks of the domes and further away: 1) the re-exposed flat sub-cambrian peneplain, 2) the re-exposed hilly sub-mesozoic peneplain and 3) epigene (never covered) plains with residual hills; i.e. the south småland peneplain (ssp), the muddus plains and the palaeic surfaces of the southern scandes (not marked). the re-exposed peneplains are tilted and often cut into different blocks, while the epigene peneplains are near-horizontal. mixed relief includes stepped plains interfering with hilly relief inside the bothnian coast in the north (mixed relief a) and sub-cambrian summit plateaux interfering with joint aligned valleys in south west finland/south central sweden (mixed relief b). jurassic outliers on the norwegian margin are indicated (from south to north bjorøy, beitstadfjorden and andøya; bøe et al. 2010). v: vätter graben; ö: östersund low area. modified from lidmar-bergström (1999) and lidmar-bergström et al. (2013). 21 dome was cut by erosion to a new base level after uplift, tilting and removal of the palaeozoic cover. this characteristic hilly relief continues northwards along the west coast of sweden and continues along the south-eastern slope of the southern scandes (fig. 7; lidmar-bergström et al. 2013). the preservation of this surface can only be explained by protection beneath a long-lasting late cretaceous and younger cover directly on basement (elvhage & lidmar-bergström 1987; lidmar-bergström 1995). the possibility of such a cover has been discussed in thermochronology studies of the area (cederbom et al. 2000; huigen & andriessen 2004). the hilly relief is present on the southern and western flanks of the south swedish dome only up to about 125 m a.s.l., where it is cut off by a horizontal surface, the south småland peneplain (figs 7, 8, 11). this surface is characterised by plains with scattered residual hills (fig. 12) and remnants of gravelly saprolites up to over 10 m thick (lidmarbergström et al. 1997). s s p s c p smp smp s m p sub-mesozoic peneplain (smp) hilly relief w e south småland peneplain (ssp) 200 m level 200 m level ssp escarpment sub-cambrian peneplain (scp), flat relief öland cambro-silurian 200 km10050 1500 granitepzgneiss 300 56°n 57°n 13°e 15°e 17°e 100 200 0 el ev at io n (m a .s. l.) a a' a a' 0 50 100 150 200 250 300 350 elevation (m a.s.l.) n pzgneiss granite ö lan d 100 km denmark s t z basement photo location palaeozoic mesozoic and cenozoic s o u t h s w e d i s h d o m e (ssd) a'a fig. 12 fig. 9a fig. 9bfig. 10b fig. 10a fig. 8. elevation of southern sweden. the south småland peneplain (ssp), an inselberg plain (fig. 12) at the south-western flank of the south swedish dome (ssd), cuts across gneiss in the west and granite/porphyries in the east. note the tilted cut-off, re-exposed sub-mesozoic peneplain (smp) (hilly etch surface; fig. 10) in the west and south and the tilted and uplifted re-exposed, flat sub-cambrian peneplain (scp) (fig. 9) in the east (profile from lidmar-bergström 1988). pz: protogin zone. stz: sorgenfrei–tornquist zone. the three different relief types are not related to differences in geology but to denudation history and can be used to reconstruct the tectonic development (lidmar-bergström 1993, 1994; japsen et al. 2002). 2222 3.1.3 cross-cutting relationships and their implications the cross-cutting relationships described above lead to the following insights. regional unconformities with remnants of subaerial weathering provide evidence of peneplain formation in the geological past. saprolites and relief (flat, hilly, etc.) are different at the contact with different cover rocks. the crosscutting relationships between peneplains of different tilts and with different relief and saprolite give information about the relative denudation chronology. a re-exposed peneplain in basement rocks can be identified at its contact with the cover rocks. its characteristic topography, saprolite remnants (lidmar-bergström et al. 1997; bonow 2005) and inclination can be followed from below its cover. where the land surface changes character (relief and saprolites) and inclination, this indicates that a former continuation of the re-exposed surface has been erased by later erosion (fig. 11). the landforms at the south-western border of the south swedish dome (fig. 11) show how such relationships can be used for extracting the geomorphological development in three stages: 1. the sub-cambrian peneplain was re-exposed to denudation along the south-western part of the south swedish dome in the mesozoic. thus a new surface formed and became graded to sea level by the fluvial system. due to a warm and humid climate an irregular relief with thick kaolinitic saprolites along fracture zones formed (fig. 10). this deeply weathered new peneplain was successively tilted, partly stripped of its saprolite and subsequently covered by sediments during the late mesozoic. 2. after a transition from subsidence to uplift, the south swedish dome rose above sea level and a younger, epigene, post-cretaceous, sub-horizontal surface (the south småland peneplain) was formed by erosion cutting across both basement of different lithologies (stephens et al. 1994) and cover rocks, a b fig. 9. a: view over the inclined subcambrian peneplain from aboda klint, 30 km west of remaining cover rocks. the sub-cambrian peneplain is an impressive regional feature. b: sub-cambrian peneplain close to its cambrian cover rocks, where it is everywhere extremely flat (fågelmara). photo locations shown in fig. 8. photo: m. rowberry, academy of sciences of the czech republic 23 guided by a new distinct base level (fig. 11b). this base level was unrelated to resistant rock types. as the cenozoic record in denmark (rasmussen et al. 2008) shows that the area was close to the sea in the west, sea level provides the only feasible base level. erosion resulting in formation of this younger surface has thus removed both the cover and basement rocks above its present level. the new plain is subhorizontal, with only a few residual hills, and exhibits gravelly saprolites. such saprolites started to form in the miocene (migoń & lidmar-bergström 2001) providing support for the argument that the plain formed during the neogene. 3. uplift/lowering of base level (here only about 100 m) has caused re-exposure of the sub-cretaceous hilly relief with associated clayey saprolites at lower elevations to the south and west of the dome. the lowering of the base level is of a late age, as the hilly relief and often also the associated kaolinitic clayey saprolites remain on the tilted sub-cretaceous surface. the present elevation of the sub-horizontal south småland peneplain indicates uplift relative to the present-day sea level. afig. 10. the precambrian basement in south sweden is weathered along zones of variable width to a kaolinitic saprolite, up to 60 m thick, below or in close connection to preserved upper cretaceous strata. a: in a quarry on the island of ivön, partial evacuation of the saprolite has exposed the steep weathering front. the irregular relief with steep-sided hills has formed by deep weathering in the mesozoic and subsequent stripping of the saprolite. note person for scale (red circle). b: typical steep-sided hill of the subcretaceous hilly relief at dalhejaberg with kaolinitic saprolite remnants in fractures. note the complete contrast to the landscape of the extremely flat sub-cambrian peneplain. photo location shown in fig. 8. b 2424 3.1.4 peneplains as unconformities in other areas and their crosscutting relationships peneplains as unconformities in the stratigraphic record, such as the sub-cambrian and sub-mesozoic surfaces of the baltic shield, are of common occurrence in many regions of the earth. for instance in germany the peneplanation of the variscan mountains was terminated with the formation of ‘der permische rumpf ’ (the permian base levelled plain), on which the upper permian and mesozoic sediments were deposited (gellert 1958). the old peneplain is re-exposed where the cover sediments have been removed by erosion. after uplift, tilting and re-exposure of the surface, younger, more horizontal peneplains have formed and are covered for instance by palaeogene sedimentary strata (baulig 1928; büdel 1977; demoulin 2003). in canada sub-ordovician flat surfaces are an important component of the present relief (ambrose 1964; lidmar-bergström & jansson 2005). 3.1.5 sea level, a major base level an extensive, low-relief erosion surface which cuts across rocks of different resistance provides evidence that its formation was governed by a particular base level for an extended time. the above examples of unconformities developed over large areas and subsequently covered by marine deposits demonstrate that the sea level acted as base level for the formation of extensive peneplains. it is not possible to argue that such low-relief unconformities which have been covered by marine strata, such as the sub-cambrian flat peneplain or the sub-cretaceous hilly peneplain, were formed at high elevations (fig. 11), as they were both progressively transgressed by the sea. the near-horizontal south småland peneplain on the south-western flank of the south swedish dome cuts off a re-exposed, inclined, sub-cretaceous surface. this latter surface would not have retained its characteristic relief and saprolites without a protective cover. thus the south småland peneplain cannot have formed at its present elevation but is part of a larger surface that originally formed across both basement (the present south småland peneplain) and cretaceous cover rocks (protecting the hilly relief) guided by a common base level (fig. 11). the south småland peneplain must have been uplifted in relation to the present base level. if continuously exposed, the relief at low elevations would have been subjected to the same weathering and erosion as at higher elevations and, in accordance with the prevailing climate, only gravelly saprolites would have formed. instead, hilly relief and kaolinisation at low elevations are observed in the west and south, where the relief disappears below a mesozoic cover. where the relief disappears below a cambrian cover in eastern sweden, the bedrock surface is flat. a low base level would have caused erosion of the re-exposed surfaces and erased them. where a study area is known to have been close to the sea during formation of a peneplain, as in south sweden (section 3.1.3) or in the case of the post-rift development of margins adjacent to opening oceans, the most likely base level is sea level (see japsen et al. 2009). thus the importance of sea level as a general base level, as advocated in the early studies (section 2.1; e.g. davis 1899; king 1962), remains valid. 3.1.6 other implications of re-exposed peneplains the once-horizontal, sub-cambrian peneplain of the south swedish dome has been uplifted, tilted and partly re-exposed long after its formation. it has been incised by valleys and is now being destroyed (fig. 8). this information leads to the conclusion that uplift leads to incision. peneplains must also have been close to horizontal when they were being formed, as tilted plains start to be dissected at a slope of 0.4–0.5% (rudberg 1970; spönemann 1979) and thus cannot survive for long after tilting. thus, they are transient features (bishop 2007). re-exposed peneplains commonly occur as tilted surfaces at low elevations, but at higher elevations such surfaces are overprinted by younger peneplains (fig. 11; garzon et al. 1982; lidmar-bergström 1988, 1996; jutras & schroeder 1999; jutras & prichonnet 2004; peulvast & claudino-sales 2004; bonow et al. 2006a; peulvast et al. 2008). well-preserved re-exposed peneplains are evidence of long-lasting covers. re-exposed peneplains define both periods of uplift/exposure and periods of subsidence/covering. 3.1.7 peneplain formation – discussion erosion by coastal waves can be responsible for the final modification of low-relief erosion surfaces during major transgressions (e.g. king 1963), but this cannot be the 25 cause of denudation of hundreds of metres or kilometres of rocks across areas of regional extent (adams 1975). the re-exposed peneplains also reveal that their final shape depends on the environmental conditions during their formation. the south sweden example shows this well with the different forms and saprolites of the three distinct surfaces: the sub-cambrian flat peneplain with shallow (maximum 5 m) kaolinitic saprolites, the hilly sub-mesozoic etch surfaces with thick (60 m) kaolinitic saprolites, and the south småland peneplain with scattered residual hills and gravelly saprolites of moderate thickness (up to about 10 m). weathering processes in combination with slope processes and surface wash tend to flatten hilly relief in arid to semiarid climates in both hot and cold environments (section 3.2.2), but cannot provide a specific base level for a plain of regional extent. the work of brozovic et al. (1997) has been used in recent papers to argue for peneplain formation at high elevations (egholm et al. 2009; steer et al. 2012). we note that brozovic et al. (1997) speculated (their wording) about a landscape between 4000 and 5000 m in the north-western himalaya with undissected plateaux, partially dissected plateaux or deeply dissected mountainous regions with comparatively gentle slopes (25 degrees), which might (our italics) be interpreted as an erosion surface. however, the relief described by brozovic et al. (1997) does not mimic what we describe as a peneplain, which, by definition, cannot have a relief of 1000 m. brozovic et al. (1997) further speculate that this region at elevations between 4000 and 5000 m is affected by both glacial scour and vigorous freeze-thaw action. this is probable, but the area is not a regional erosion surface or peneplain. mitchell & montgomery (2006) showed that the distinct peak accordance with a summit plane cannot a priori be regarded as remnants of a former peneplain. they used the ‘glacialbuzzsaw’ mechanism, proposed by meigs & sauber (2000) to show that mountains in the cascade range, western usa, have been cut off at a certain elevation averaging 380 m and maximally 600 m higher than the equilibrium-line altitude (ela), visualised as an inclined plane defined as the floors of 373 cirques in areas exposed to abundant precipitation. on the other hand, the peneplains described here from southern sweden are not dissected summit surfaces but coherent plains that require another mechanism than the ‘glacial buzzsaw’ for their formation. mitchell & montgomery (2006) state that the ‘glacial-buzzsaw’ mechanism includes three base level (sea) base level (sea) epigene peneplainre-exposed peneplain a. present. after lowering of base level and/or uplift deposition formation of epigene peneplain across basement rocks b. earlier stage. after uplift, tilting and erosion covered peneplain (unconformity) peneplain types in basement etch surface (hilly relief) with kaolinitic (ka) saprolite inselberg plains with shallow or no saprolite basement mesozoic and younger sediments kk cretaceous outlier kk covered peneplain (unconformity) change of: 1) surface inclination 2) relief type 3) saprolite type ka ka formation of epigene peneplain across cover rocks erosion fig. 11. cross-cutting relationships between peneplains. a: the relationship between a tilted, re-exposed peneplain and a subhorizontal epigene peneplain, each with its own characteristic relief and saprolite, determines their relative ages. the subhorizontal peneplain, cutting off the tilted peneplain, is the youngest. this case is shown at the western border of the south swedish dome, where a mesozoic peneplain of etch character is cut off by a cenozoic peneplain with scattered inselbergs (lidmar-bergström 1988). b: the landform record can be used to extract landscape history in a tectonic/eustatic context. in an earlier stage, before lowering of the base level and/or uplift, the epigene peneplain formed across both basement and cover rocks by grading to a former base level. formation of the epigene peneplain has required erosion first of the former cover and second of basement rocks. this stage must have been preceded by uplift and tilting (see sections 3.1.3, 3.1.5). 2626 distinctive characteristics in relation to the quaternary ela: 1) a line through the peaks is never located more than 600 m above the surface defined from the bases of the cirques; 2) the amount of topography declines above the ela and 3) the valley sides above the ela are close to threshold steepness. thus there is no plain at all above the ela but a valley landscape with steep slopes with a relative relief of up to 600 m (fig. 13a). in the original paper on the ‘glacial buzzsaw’, meigs & sauber (2000) state that the mean height of topography and the mean height of the ela are coupled. however, this does not mean that there are low relief surfaces similar to peneplains. in the western areas studied by meigs & sauber (2000), for example, the land lies in an altitude band between 300 and 2200 m with the highest summit at 6050 m. thus there is relative relief of at least 3850 m. despite this, egholm et al. (2009) suggested that the equilibrium-line altitude (ela) that forms the base level for enhanced glacial erosion of glaciers and cirques can lead to the formation of flattish surfaces, which grow together and form extensive low-relief surfaces at a distinct level. this is not in accordance with geomorphological studies, which instead show that glacial erosion amplifies the relief (sugden 1978; lidmar-bergström 1997; johansson et al. 2001b; kessler et al. 2008) or dissects earlier formed surfaces (oskin and burbank 2005; etzelmüller 2007). it is certainly not valid as a mechanism for the formation of low-relief surfaces of regional extent, and a flattish surface of minor extent is not a peneplain. egholm et al.’s (2009) hypothesis cannot be applied to all epcms, since many of them (e.g. in eastern australia, southern africa and eastern brazil) have not been significantly modified by glaciation since palaeozoic times. steer et al. (2012, p. 1) studied low-relief surfaces at different elevations in the sognefjord catchment area of norway. they argued that these surfaces were shaped by glacial erosion at ela during different times and wrote: “these surfaces have been attributed to glacial headward erosion in alpine settings”. however, the notion of an alpine setting refers to mountainous areas with glacial forms such as cirques and glacial valleys and not to peneplains. steer et al. (2012) appear to suggest that low-relief surfaces can be formed by glacial headward erosion by cirque retreat as they refer to oskin & burbank (2005), but the process they describe simply amplifies the relief (see above), and it has not been shown anywhere that glacial valleys coalesce to produce low-relief erosion surfaces. this type of erosion cannot form the type of regional peneplains that characterise epcms all over the world, glaciated or not. in addition, the main distribution of high-level peneplains in southern norway does not occur on its western and north-western side with its alpine relief (fig. 4), which developed where precipitation was highest and the frequency of cirques largest. instead, the best developed high-level plains with residual massifs occur in the east, least affected by glacial erosion (sollid & sørbel 1994; kleman et al. 2008). further evidence that glacial erosion acts to dissect plateaux rather than create them in western norway was recently published by hall et al. (2013). it should also be noted that the regional and tilted upper planation surface (ups) in west greenland, which covers a region more than 300 km long and 145 km wide between the coast and the greenland ice sheet and reaches heights over 1500 m a.s.l. in the south-west and 200-500 m a.s.l. in the north and east (bonow et al. fig. 12. plain with residual hill on the south småland peneplain, the end result of cenozoic landscape formation. photo location shown in fig. 8. 27 2006a), cannot be correlated with any ela (see section 5.3 for further discussion on this topic). steer et al. (2012) also referred to hales & roering (2009) who described frost processes at high elevation in new zealand that prevent peaks rising above 2300 m. hales & roering (2009) did not, however, describe the formation of peneplains at high elevation but instead suggested that frost processes cause rock falls, in turn causing the summits to be at the same elevation despite continuous uplift. high plains, in general, are coherent features. they are best preserved in non-glaciated areas where they cannot have been shaped by glacial processes. peneplains at high elevations can still be recognised even on glaciated margins (fig. 13; bonow et al. 2003, 2006a, b), despite glacial reshaping. no credible mechanism has as yet been suggested to produce regional peneplains at high elevations, unless controlled by a structural/lithological base level, but their formation close to sea level is demonstrated by both re-exposed peneplains and their cross-cutting relationships to younger peneplains in positions close to former seas, as has been described in section 3.1.5 and preceding sections. in conclusion: low-relief erosion surfaces extending over hundreds of kilometres, identified as re-exposed peneplains, provide evidence of peneplain formation close to sea level and are also important indicators of former covers. their crosscutting relationships to higher surfaces also confirm the sea as base level for these surfaces that are now at higher elevations. 3.2 terminology 3.2.1 landscape, relief and topography both the words ‘landscape’ and ‘relief ’ are used to describe landscapes with slopes of different character, landforms characteristic of different processes in different environments and regions with different relative heights. in sla we use ’relief ’ to mean the relative variation in height of the landscape, and the word ‘topography’ to mean the absolute heights of the landscape above sea level. 3.2.2 peneplanation or pediplanation the development of low-relief erosion surfaces graded to distinct base levels by fluvial incision and slope decline/retreat, is fundamental in the landscape models of both davis (1899) and king (1962; 1967). the main difference between them is how they defined slope behaviour. while davis based his model on slope decline, king regarded parallel scarp retreat along incised valleys as the fundamental process (fig. 14). king labelled the resulting low-relief erosion surface a ‘pediplain’. the exact behaviour of slope development is now known to depend on details in lithology, structure and fracture systems in combination with different climates (moon & selby 1983; gunnell 1998), and deep weathering plays an important role in the development of slopes between different surfaces (peulvast 1987; gunnell 1998). an example of the formation of a minor scarp is where the south småland peneplain of south sweden cuts back into the uplifted sub-cambrian peneplain (fig. 8; lidmar-bergström 1988; olvmo et al. 2005). resistant rocks can give rise to prominent escarpments such as the drakensberg escarpment (fig. 2; gunnell 1998; fjellanger & etzelmüller 2003), because they cause backward erosion to slow down (fig. 14). river gaps can form across resistant rocks by superimposition of drainage. the related surface will continue to form behind the obstacle subsequent to breakthrough, and surface development will be more rapid in the less resistant rock. the nature of the slope between two surfaces, either an escarpment or gradual slope, is a function of rock resistance and climate. in conclusion: ‘peneplain’ should be used as the general term for extensive, low-relief erosion surfaces, hilly or flat, graded to a distinct base level. for the purpose of using peneplains as markers of uplift the exact nature of slope development is not critical. the focus must instead be on the level of the erosional base for the peneplain. 3.2.3 pediplains (rock-cut plains), pediments, etched surfaces (hilly relief), inselberg plains and etchplains the term ‘pediplain’ is also used as a descriptive term, in the sense of a rock-cut plain (büdel 1970). flat, minor rock-cut surfaces with shallow or no saprolites are often termed pediments (e.g. dohrenwend 1994; vincent & sadah 1996). they have low slope angles, <6°–11°, ac2828 fig. 13. comparison of a ‘glacial-buzzsaw’ landscape with landscapes typical of elevated, passive continental margins (epcms). 3d pictures of different mountainous landscapes taken from google earth (no vertical exaggeration). all three pictures were taken from the same virtual height of c. 6.8 km. a: cascade mountains, usa. typical mountain range affected by the ‘glacial buzzsaw’ with cirque floors (cf) that constitute a tilted planar zone at the equilibrium-line altitude (ela; mitchell & montgomery 2006). above the cirque floors is an alpine landscape with a relief between 380 and 600 m, but there is no plain despite the claims of egholm et al. (2009). this type of alpine landscape contrasts dramatically with the high-level epcm peneplains shown in b and c. b: norwegian epcm. hardangervidda, looking approximately south. an elevated peneplain (c. 1200 m a.s.l.) is level 3 in the classification of lidmar-bergström et al. (2000). in contrast to the ‘glacialbuzzsaw’ situation in a, the valleys to the right (west) have been incised below the peneplain. glacial erosion has deepened their floors to near or below present-day sea level, because the pleistocene ela in norway was below sea level. c: eastern australian epcm. plateau in kosciusko national park, looking slightly east of north. unglaciated epcm with an uplifted peneplain to the left (west); 1400–1600 m a.s.l. fluvial valleys (now flooded by a hydroelectric lake) are eroding into the peneplain from the right (east). level 3 surface ca. 1200 m a.s.l. sea uplifted peneplain 1400–1600 m a.s.l. river incision cfcf lakes at 1150 m a.s.l. a b c cfcf cfcf cfcf cfcf cfcf 29 cording to dohrenwend (1994) or <6.5° according to bonow et al. (2003). the higher figure seems to denote the slope in direct contact with residual hills. pediments can grow large, forming pediplains by scarp retreat in arid climates around residual hills, exposed after erosion of an originally deep saprolite and resulting in elimination of all but a few hills (thomas 1974; demangeot 1976; lidmar-bergström 1982, 1995) or as pedimentation along the sides of valleys and eventual removal of intervening ridges (ahnert 1998). pedimented surfaces evolve in arid and semiarid zones by slope retreat and surface wash (büdel 1970; mensching 1970; young 1972). the sub-cambrian peneplain in fennoscandia is extremely flat, only a little weathered (see above) and with very few residual hills. it was formed at a time without any vegetation cover, when continuous surface wash therefore contributed to its final shaping as a pediplain (lidmarbergström 1988). both shallow weathering and surface wash are ingredients in pediplain formation. on the other hand, field observations in the semihumid tropics have shown that differential weathering by long periods of deep weathering with formation of deep saprolites (up to over 100 m) alternating with periods of stripping of the saprolite, followed by renewed deep weathering (thomas 1965, 1966) result in a highly irregular weathering front. after final exposure of the weathering front this gives rise to more or less closely spaced hills, depending on fracture density. in fennoscandia and greenland, lidmar-bergström (1982, 1989, 1995, 1999) and bonow (2005) have shown that hilly relief (irregular bedrock surfaces with hills up to 200 m and more) was caused by periods of kaolinitic weathering (etching) and stripping in subtropical climates in the mesozoic (fig. 10). such surfaces are only preserved where they have been protected until recently by mesozoic–palaeogene covers (fig. 11). re-exposed hilly relief is often reinforced by glacial erosion in the formerly glaciated areas (johansson et al. 2001a; migoń & lidmar-bergström 2001; bonow 2005), while it can be partly hidden below thick saprolites in tropical and subtropical areas. the end product of relief formation in greenland and scandinavia during the cenozoic is plains with residual hills with thin, gravelly, saprolites (lidmarbergström 1982; elvhage & lidmar-bergström 1987 1995; lidmar-bergström et al. 1997; bonow et al. 2006a, b; lidmar-bergström et al. 2007). in contrast to the subcambrian peneplain, which is almost devoid of residual hills, and to the hilly sub-mesozoic etched surfaces, the cenozoic plains are characterised by isolated residual hills (inselbergs; fig. 12) and are termed inselberg plains (rudberg 1960, 1988; ebert et al. 2012). the observed cross-cutting relationships of peneplains of different shape suggest that the inselberg plains in the tropics are produced in a similar way, rather than by deep weathering to an even, horizontal weathering front as is often suggested (e.g. moore et al. 2009). the irregular weathering guided by joints and fractures instead seems to have originally caused hilly relief (see above) from which the present rock-cut plains with only a few inselbergs have developed (lidmar-bergström 1995). etch planation is not regarded as a general process for the formation of low-relief surfaces unrelated to base level as suggested by thomas (1994). instead we define an etch plain, as a plain underlain by a deep saprolite but which also formed in relation to a distinct base level. in conclusion: pediplains, inselberg plains, etched surfaces (hilly relief) and etch plains are the result of different climatic conditions during the final stage of peneplain development. however, the formation of the peneplain as such, which involves the kilometre-scale erosion of rock during valley incision and valley widening by running water and slope processes, is guided by a stable base level, often the sea (davis 1899). 3.2.4 classification of peneplains the notion ‘planation surface’ objectively describes a flat landscape form (adams 1975). not all peneplains are planation surfaces. a hilly relief surface caused by irregular deep weathering and belonging to an identifiable plain governed by a base level is a peneplain. there is a plethora of different terms in common use and much confusion surrounds the terminology (ebert 2009). a classification is listed in table 1. a peneplain is a lowrelief erosional surface graded to a distinct base level. it can be either underlain by a deep saprolite (etch plain) or cut across fresh bedrock, almost without inselbergs (pediplain) or with scattered inselbergs (inselberg plain). etched bedrock surfaces can be more or less stripped giving rise to a hilly relief, particularly common in formerly glaciated basement terrain where stripping has been efficient. empirical figures for relative relief are available from glaciated fennoscandia (lidmar-bergström 1995); pediplains <20 m, inselberg plains c. 20–200 m (plain between scattered hills <20 m) and hilly relief c. 20–200 m. 3030 3.2.5 high-level benchlands (stepped peneplains), incised valleys and the great escarpment there is a clear difference between pedimented surfaces with low slope angles (section 3.2.2) and steeper slopes in the landscape. many areas with high-level plains and residual mountain massifs are characterised by benchlands of pediment character flanking the individual massifs with intervening steeper slopes. besides along some continental margins as e.g. in southern africa (fig. 2), such benchlands characterise e.g. kenya (ahnert 1982), spain (casa-sainz & cortès-gracia 2002) and southern norway (bonow et al. 2003). the benches sometimes occur along major river valleys and may indicate former base levels for the fluvial systems and related peneplains (bonow et al. 2003). the valleys may have coalesced to partial peneplains which, after step-wise lowering of base level, now fringe the massifs as benchlands. there is a major difference between high-level peneplains/benchlands with their gentle relief down to a certain base level plain and valleys deeply incised from this lowest plain. the zone with incised valleys along epcms is often named the great escarpment. uplifted peneplains with benchlands are called palaeic surfaces in norway (reusch 1901) and relict surfaces in southeastern tibet (schoenbohm et al. 2004; clark et al. 2006). the once horizontal, sub-cambrian peneplain of the south swedish dome in south sweden is uplifted, tilted, re-exposed and incised by valleys (section 3.1.6), which shows that uplift causes incision. the main base for the lowest high planation surface, from which deep valleys are incised, is regarded as a major marker for calculations of amounts of late uplift (bonow et al. 2006a, b). the cross-cutting relationships of such a surface to a tilted re-exposed peneplain can attest this conclusion (section 3.1.3). major knickpoints in river profiles (fig. 14) might be markers of former base levels (schoenbom et al. 2004; crosby & whipple 2006; bridgland & westaway 2008). steep channel profiles incised along the front of the tibetan plateau have led to the conclusion that such steep profiles can reflect active uplift of the plateau (kirby et al. 2003). in glaciated areas only knickpoints reflecting incision in major peneplains can be used for deciphering uplift events, while knickpoints in incised rivers might just reflect uneven glacial erosion. knickpoints can also form in tributary valleys due to glacial deepening of major valleys (kleman & stroeven 1997; bonow et al. 2003) and are not usable as uplift markers in such settings. 3.3 palaeoplains, preservation, destruction and age when a peneplain is buried beneath a sedimentary or volcanic cover, or loses its direct contact with its base level due to uplift, it is a palaeosurface (bonow et al. 2006a) and can be labelled a palaeoplain. cover rocks can preserve palaeoplains in the form of unconformities for long periods of time. epigene peneplains that have lost contact with their general base level continue to grade to their own uplifted base level and saprolites can grow (e.g. carmo & vasconcelos 2004). alternatively, the high surfaces can be fossilised by extreme aridity (migoń & goudie 2001) or below cold-based ice frozen to the ground (sugden 1968; dyke 1993; kleman 1994). areas that have experienced continued extreme aridity such as namibia and the dry valleys of antarctica (cockburn et al. 1999; summerfield et al. 1999) might contain very old features, which make it possible to reveal a history comparatively far back in time. different types of duricrusts, such as ferricrete, silcrete and calcrete, can form and will then make weathered surfaces resistant to destruction. once a peneplain is formed and uplifted (a palaeoplain), it will survive much longer in resistant rock than in more easily eroded rocks, where a new generation of valleys will first be formed (gunnell 1998; bonow et al. 2003, 2006a; fjellanger & sørbel 2007). narrow planation surface pediplain relative relief <20 m† inselberg plain relative relief c. 20–200 m etchplain hilly relief etched hilly relief relative relief c. 20–200 m after total stripping in glaciated terrain rock-cut, flat surface surface with scattered residual hills flat surface underlain by deep saprolite irregular bedrock surface with closely spaced hills caused by differential deep weathering and stripping of saprolite table 1. classification of peneplains§ in basement terrain § peneplains: low-relief erosional surfaces graded to a distinct base level. † figures for relative relief from rudberg (1960). 31 valleys are the first sign of destruction. the dissection of the south-dipping basement plain of the recently uplifted shillong plateau in india is a good example (fig. 15; cf. fig. 3; biswas & grasemann 2005; biswas et al. 2007). valleys can only incise due to a lowered base level and elevated surfaces are destroyed successively by fluvial incision and subsequent valley widening (king 1967; ahnert 1982, 1998; gunnell 1998; lidmar-bergström et al. 2000, 2007; bonow et al. 2003, 2007a). stroeven et al. (2009) noted the effectiveness of the fluvial system for dissection of the north-east tibetan plateau. high-level peneplains are often assigned very old ages even far away from the contact between the basement and its cover rock (e.g. twidale 2007). it is argued that they must be older than the cover but without checking if it really is the same surface as at the cover contact. we have noted that such peneplains in fennoscandia and greenland always cut off the adjacent re-exposed surface and are thus younger (fig. 11). they are not continuations of any re-exposed surface, but new surfaces. thus the old surface has been destroyed and a new surface has formed (sections 3.1.2, 3.1.3). experience in greenland (japsen et al. 2006) suggests that epigene surfaces older than miocene are unlikely to be preserved unless they have been carved in extremely resistant rocks or preserved in a landscape of high aridity (cf. section 8.1). re-exposed surfaces can be very old and they are easy to date approximately. in conclusion: a palaeoplain only exists for a certain time, which varies with different climatologic, tectonic and lithological settings. palaeoplains (uplifted epigene or re-exposed surfaces) are therefore transient features (cf. bishop 2007). fig. 14. generalised model of relief development based on landscape analysis of southern norway and southern sweden. major surface remnants (i) on both resistant and non-resistant rocks as a continuation of a common plane are indicative of a dissected higher peneplain. slope decline according to davis (1899) and parallel scarp retreat according to king (1962, 1967) is illustrated. the mode of valley incision to a new base level is now known to be dependent on climate and rock type. superposition of fluvial valleys from the higher base-level plain produces river gaps through resistant rocks and continued surface formation behind the obstacle. resistant rocks can give rise to temporary escarpments and residual hills. note that at low elevations valleys (iii and iv) might be incised in re-exposed relief. the near-horizontal peneplain ii cuts off a re-exposed and inclined sub-mesozoic surface with hilly relief. this shows that peneplain ii formed after uplift following deposition of the mesozoic sediments, e.g. during the cenozoic as in the case of the south swedish dome where remnants of the upper cretaceous cover are preserved on the hilly relief at the flanks below the south småland peneplain (see fig. 11). parallel scarp retreat knickpoint knickpoint slope decline re-exposed, tilted peneplain with hilly-relief sea epigene peneplains mesozoic outlier temporary local knick-point valley generations mesozoic basement resistan t basement 3232 3.3.1 effects of glaciation glacial erosion can aid in re-exposing palaeosurfaces as it easily erodes cover rocks, particularly those that are not well consolidated. glacial erosion easily evacuates saprolites, and therefore hilly etch surfaces are accentuated in formerly glaciated terrain, e.g. the sub-mesozoic relief in south sweden (olvmo et al. 1999). how much saprolite remains in formerly glaciated terrain depends on the degree of glacial erosion (hall & sugden 1987; lidmar-bergström 1997). on the other hand regional flat peneplains remain flat after glacial erosion even in regions with so-called scoured bedrock (johansson et al. 2001b). scoured refers to the occurrence of thin patchy drift and the grinding of the bedrock surface (kleman et al. 2008) but not to deep erosion of hard, fresh rock. in general, glacial erosion reinforces preglacial relief by preferentially eroding in valley floors and valley sides (ljungner 1949; sugden 1978; rudberg 1992; fredin 2002, bonow et al. 2003) rather than on summits, which were protected from erosion by the ice cover due to coldbased conditions over long periods of time (sugden 1978; kleman 1994; kleman & stroeven 1997; kleman & glasser 2007; kleman et al. 2008). the deepest glacial erosion is found along major valleys that have hosted outlet glaciers from large ice sheets, where glacial erosion to over 1000 m below sea level has been documented, for example in sognefjord, norway (see fig. 4; nesje & whillans 1994). knickpoints are destroyed but remnants of earlier valley bottoms may occur as local valley benches (bonow et al. 2003). there is also a significant difference in the effect of glacial erosion over basement terrain and across sedimentary rocks as shown by the substantial erosion along the norwegian coast (rise et al. 2005). in mountainous terrain the erosion of valley sides and valley floors by valley glaciers makes the pattern of preglacial slope and knickpoint development difficult to reconstruct (cf. bonow et al. 2003). cirques destroy palaeosurfaces to form an ‘alpine relief ’ (figs 4, 5), where 92°e91.5°e91°e 25°n 26°n 25.5°n 10 km n0 500 1000 1500 2000 2500 elevation (m a.s.l.) india fig. 15. shillong plateau, india. the precambrian basement was exposed in the late cretaceous and subsequently buried and covered by thick cenozoic deposits. the gently south-dipping basement plain is re-exposed from these strata after pliocene uplift (biswas et al. 2007). the southward-facing, inclined peneplain is currently being dissected by deep valleys down to the new base level, similar to the development of the east australian escarpment and coastal plain (fig. 3). the succeeding valley-widening process, ultimately resulting in a new base-level-governed peneplain, depends on lithology and weathering conditions. 3d-terrain model with vertical shading. elevation data from jarvis et al. (2008). 33 elevations are sufficiently high and the climate is suitable (etzelmüller et al. 2007). as discussed in section 3.1.7, a ‘glacial-buzzsaw’ effect can reduce mountain heights in areas exposed to high precipitation and can produce summits of equal heights (mitchell & montgomery 2006), and the equilibrium line altitude (ela) can form a base level for intensified glacial erosion in valleys (brozovic et al. 1997). on the nuussuaq peninsula and the island of disko in central west greenland, the upper planation surface (ups) is tilted to the west and well preserved below 1000 m a.s.l. (see chapter 5), while it is more or less destroyed above this level close to the coast (bonow et al. 2006b). the ups is again identifiable 100 km eastwards, where it is a coherent surface descending from 2000 m a.s.l. in the west to 500 m a.s.l. in the east (bonow et al. 2006b). in the intervening region, glacial erosion has limited the summit heights to about 1500 m a.s.l., possibly by the ‘glacial-buzzsaw’ mechanism (mitchell & montgomery 2006). in conclusion: peneplains in stepped sequences can also be identified in formerly glaciated areas, while preglacial valley steps are destroyed. in areas close to a coast, high peneplains might be dissected by cirques and valley glaciers. 3.4 combining sla and thermo chronology development of peneplains clearly requires erosion of large amounts of rock. analysis of peneplains, re-exposed and epigene, and their cross-cutting relationships, provides basic information on the relative chronology of their formation, subsidence, burial, uplift, tilting and formation of new peneplains. the most recent uplift relative to sea level is revealed by re-exposure of formerly covered peneplains and also by valley incision below the lowest of uplifted peneplains. sla is based on integration of information on the character of surfaces (forms and saprolites), the relationships between them, and geological constraints. special emphasis is placed on identifying sedimentary remnants which define re-exposed peneplains. where applicable, the effects of glacial reshaping should also be taken into account. sla is independent of thermochronology so the two disciplines are discussed separately here. the relative chronology of events provided by sla can be placed within an absolute frame of reference using low-temperature thermochronology. integration of the two approaches allows determination not only of the timing of erosional episodes but also quantifies amounts of removed section, including both basement and sedimentary cover rocks. this is discussed in the next chapter. 3434 4. low-temperature thermochronology 4.1 apatite fission-track methods 4.1.1 historical background naturally occurring fission tracks are radiation-damage trails produced by the spontaneous fission of 238u atoms, in which a uranium atom splits into two fragments. these fragments are stripped of electrons and positively charged, so they repel each other through the lattice and create a linear damage zone consisting of displaced atoms. these damage zones are highly chemically reactive, and can be selectively dissolved and thereby enlarged by a simple etching treatment. thus fission tracks produced by spontaneous fission of individual u atoms within an apatite crystal can be revealed where they intersect a surface (fig. 16). the number of tracks per unit area of a polished and etched grain surface is controlled by uranium content and time, through standard decay laws, and track length, through geometrical considerations (galbraith 2005). therefore, by counting the number of tracks and measuring the uranium content, a ‘fission-track age’ can be calculated which, in the absence of other factors, should indicate the time over which tracks have accumulated. the uranium content is measured by irradiating a grain mount with thermal neutrons, which produces induced fission of 235u atoms in the apatite grains. some of the resulting fission fragments are emitted from the grain surfaces and are recorded as ‘induced tracks’ in a muscovite external detector after etching. the ratio of spontaneous tracks (daughter product) to induced tracks (parent) is converted to a numerical age using age standards (e.g. hurford & green 1983; green 1985). reviews of the basics of fission-track dating are provided by e.g. fleischer et al. (1975), wagner & van den haute (1992) and galbraith (2005). early applications of apatite fission-track dating to accessory apatite grains extracted from granitic rocks suggested that fission-track ages could be reset at relatively low temperatures around 100°c over geological timescales (e.g. wagner & reimer 1972). this observation was supported by early laboratory annealing studies (wagner 1968; naeser & faul 1969), and subsequently confirmed by direct measurement of fission-track ages in subsurface samples (naeser & forbes 1976). integration of fission-track ages with ‘confined track’ length measurements (fig. 17), first reported by bhandari et al. (1971), led to a deeper understanding of the method. confined tracks are totally enclosed within the body of the crystal and are revealed when the etchant penetrates cracks or other tracks and intersects tracks below the surface of the grain. for this reason, the entire length of the track can be etched and measured, in contrast to the tracks used in age determination, which are truncated by intersection with the surface. early measurements showed that whereas new tracks produced by induced fission of 235u were characterised by mean confined track lengths around 16 µm, spontaneous tracks are always shorter. even in volcanic rocks which have only experienced very low temperatures after initial post-eruption cooling, mean confined track lengths are typically around 14 to 15 µm (gleadow et al. 1986a). for both types of track, the distribution of track lengths showed a standard deviation of c. 1µm, due to the variation in mass and energies of the fission fragments. green (1980) showed that the track shortening by 1–2 µm in these volcanic apatites can be explained in terms of thermal annealing of these tracks at low temperatures (<50°c) over geological timescales, highlighting the sensitivity of the technique. donelick et al. (1990) later showed that a small degree of initial track-length reducfig. 16. spontaneous fission tracks in an apatite grain, revealed after etching for 20 seconds in dilute nitric acid. the track openings are 1–2 µm in width. individual tracks range up to a maximum of c. 15 µm in length. 35 tion proceeds quickly even at room temperature for fission tracks produced by induced fission, further emphasising that annealing proceeds at finite rates even at low temperatures. results from boreholes in the mesozoic–cenozoic otway basin of south-east australia (gleadow & duddy 1981) provided the first rigorous constraints on the thermal stability of tracks in geological conditions, revealing the progressive reduction in fission-track age with increasing depth and temperature, and showing that this was complemented by a corresponding decrease in mean confined track length (fig. 18). laboratory studies (green et al. 1985, 1986; green 1988), together with detailed mathematical analysis (laslett et al. 1982; galbraith & laslett 1988, 1990), established that the progressive reduction in track length causes the reduction in fission-track age, by reducing the proportion of fig. 17. confined fission tracks in apatite, revealed by focusing below the grain surface. confined tracks can be revealed by intersections with other tracks (track-in-track or tints; black arrows) or with cracks or fractures (tracks-in-cleavage or tincles; white arrows). 0 2 4 6 8 10 12 14 16 temperature (°c) stratigraphic age 120 100 80 60 40 0 0 20 40 60 80 100 120 140 140 a b 20 m ea n fis sio ntr ac k le ng th (µ m ) temperature (°c) 0 20 40 60 80 100 120 140 fi ss io ntr ac k ag e (m a) fig. 18. fission-track ages and mean confined track length in apatites extracted from samples of early cretaceous volcanogenic sandstones in a number of wells located in the otway basin, se australia. the sedimentary sections intersected in these wells are characterised by simple histories involving continuous burial, such that all samples are now at their maximum postdepositional temperatures. the decrease of both fission-track age and mean track length with down-hole temperature therefore provides a direct expression of the thermal sensitivity of fission tracks in these apatites over geological timescales. modified from green & duddy (2013). 3636 tracks that can intersect a polished grain surface. the reduction in track length is, in turn, a manifestation of the reduction in the degree of damage within the track region as displaced atoms return to their original lattice sites by thermally activated diffusion. recognition of track-length reduction as the controlling process has underpinned all subsequent efforts to quantitatively predict apatite fission-track parameters and extract thermal history information from such data. 4.1.2 thermal response of fission tracks in apatite early laboratory studies of the thermal sensitivity of fission tracks in apatite (e.g. wagner 1968; naeser & faul 1969) were based on measuring the reduction of fissiontrack age resulting from various heat treatments. the large uncertainties in measured ages or densities, combined with possible unrecognised effects due to compositional differences (as discussed later) introduces considerable uncertainty into such measurements, allowing a range of interpretations. with the advent of confined track-length measurements, the greater precision and reproducibility of such measurements for representing the degree of annealing allowed significant refinement in the quantitative understanding of the kinetics of fissiontrack annealing in apatite. laslett et al. (1987) showed that the variation of mean track length with temperature and time in laboratory annealing studies could be described by a ‘fanning arrhenius plot’ model (adopted from earlier studies), in which contours of equal track-length reduction form straight lines in a plot of time against inverse absolute temperature, the slopes of which (reflecting an ‘activation energy’) increase as the degree of annealing increases. laslett et al. (1987) favoured a model in which all contours of equal annealing converge to a point at infinite temperature (1/to = 0), because it allowed a simplified mathematical treatment. a variety of alternative annealing models have subsequently been published (crowley et al. 1991; laslett & galbraith 1996; ketcham et al. 1999). all of these represent refinements or alternative forms of the basic fanning arrhenius plot model proposed by laslett et al. (1987), e.g. involving finite values of 1/to, or curved contours, and the principles described in the following sections also apply to these models. the improved definition of the kinetics of fissiontrack annealing provided by using mean confined track length as the fundamental parameter (laslett et al. 1987), combined with a detailed understanding of the way in a time (ma) 100 0.0 0.2 0.4 0.6 0.8 1.0 le ng th re du ct io n l/l o track-length distribution track-length distribution 0 5 10 frequency frequency te m pe ra tu re (° c ) 100 80 60 40 20 0 110 020406080 100 020406080 b time (ma) 100 0.0 0.2 0.4 0.6 0.8 1.0 le ng th re du ct io n l/l o 0 5 10 15 te m pe ra tu re (° c ) 100 80 60 40 20 0 110 020406080 100 020406080 fig. 19. predicted track-length reduction vs. time for tracks formed at different times within histories involving continuous cooling (a) and heating (b), together with final predicted track-length distributions (right). the trajectories and final track-length distributions are calculated following the methods outlined by green et al. (1989). whereas in the cooling history, tracks formed at different times end up at the present day with different lengths, in the heating case almost all tracks are shortened to the same degree, reflecting the dominance of temperature over time in the kinetics of annealing. track length reduction, l/l0, where l is measured mean track length and l0 is the initial (unannealed) mean track length. 37 which reduction in track length is manifested in the fission-track age (green 1988), provided the basis for the quantitative modelling of the response of fission tracks in apatite to different styles of thermal history in geological conditions (green et al. 1989). a key step in this process is the transition from isothermal annealing models to histories in which temperature varies with time. duddy et al. (1988) provided a way forward by adopting the principle of ‘equivalent time’ (originally postulated by goswami et al. 1984), by which the rate of annealing of a track at any given time only depends on the length to which the track has already been reduced and the prevailing temperature, and not on the history of how the track reached that length. duddy et al. (1988) verified this assumption in a series of variable temperature-annealing experiments. a key outcome of this work is the recognition of the dominance of temperature over time in fission-track annealing. in thermal histories involving continuous cooling, tracks produced at different times throughout the history are shortened at a rate which is initially high but is quickly reduced to a minimal level as the temperature falls through time. tracks produced at different times experience declining temperatures and are therefore shortened to different degrees (fig. 19a). in contrast, for histories involving continuous heating, because temperature dominates over time in the kinetics, the lengths of tracks produced at different times are progressively reduced through time, and all tracks are shortened by more or less the same amount, except for those produced within the last few per cent of the history (fig. 19b) although it should be appreciated that tracks are produced with a finite spread of lengths, as discussed below. these contrasting styles of behaviour have serious implications for the response of the apatite fission-track system in thermochronology. figure 20 illustrates the development of apatite fission-track parameters through a history involving heating followed by cooling and subsequent residence at low temperature. during the heating phase all tracks formed at different times are progressively reduced in length, and at the palaeothermal maximum all but a small proportion of tracks that have formed during the heating phase are shortened to the same degree. tracks formed after the onset of cooling remain longer, reflecting the lower temperature. at the present day, two populations of tracks are present, one with short lengths representing tracks formed during the heating phase, and another comprised of longer tracks formed after the onset of cooling (fig. 20). the population of shorter tracks will contribute a reduced component to the fission-track age, compared to the time interval over which tracks have been retained, while the contribution to the fission-track age from the longer population will be much closer to the time interval after the onset of cooling. the final measured fission-track age will represent the summed contributions of both components. a sample which reached a maximum palaeotemperature sufficient to reduce track lengths to zero prior to cooling will only contain a single population of tracks, formed after cooling to a temperature at which tracks are retained (analogous to the population of longer tracks in shallower samples). the fission-track age in such a sample will be determined by the time since the sample began to retain tracks, but the precise value will depend on the degree of length reduction of tracks formed during the cooling history. figure 21 illustrates how this results in a characteristic variation of apatite fission-track parameters with depth in a sedimentary section that has been heated and then cooled. 4.1.3 variation in annealing kinetics between different apatite species studies of apatite fission-track parameters in subsurface samples from the otway basin, south-east australia (gleadow & duddy 1981; green et al. 1985, 1986), showed that the chlorine content of the apatite grains exerts a systematic influence on annealing rates (fig. 22). this was subsequently confirmed in laboratory studies by carlson et al. (1999) and barbarand et al. (2003; fig. 23). although both these studies tended to downplay the influence of chlorine in favour of other factors, the importance of differential annealing within individual samples related to wt% cl has been demonstrated in a number of studies (e.g. argent et al. 2002; crowhurst et al. 2002; green et al. 2002; green 2005; green & duddy 2013). in contrast, evidence for systematic differences in annealing rates in geological samples due to any element other than cl has yet to be demonstrated. a correlation between etch pit diameters and annealing rates has been used in a number of studies to allow for differential annealing between different apatite species (ketcham et al. 1999), although green et al. (2005) showed that annealing rates correlate much more strongly with wt% cl than with etch pit size (see also green & duddy 2013). in geological conditions, differential annealing effects within individual samples are maximised in rocks which have been heated into the critical temperature range 3838 (typically 90–120°c). over this range, the most sensitive (i.e. low-cl) apatites are totally annealed while more resistant apatites (high-cl) are unaffected (cf. fig. 22). in such cases, the systematic dispersion in fission-track age, correlating with wt% cl, can provide added precision to a thermal history solution (green et al. 2002). an example of an apparently anomalous observation that can be simply explained in terms of differences in wt% cl is the difference in fission-track age between adjacent gneissic and charnockitic terrains in india, which gunnell (2000) attributes to differing resistance to erosion. this is more likely to be due to a difference in annealing rates in apatites from the two rock types, with charnockitic apatites likely to be richer in cl, and hence giving older ages. in such cases, measurement of cl content by electron microprobe can easily resolve such effects. lorencak et al. (2004) provide graphic evidence of the impact of small-scale variations in apatite cl content within a single rock body on measured apatite fission-track ages due to the influence of wt% cl on annealing rates. sample 2 le ng th re du ct io n l/l o 1.0 0.8 0.6 0.4 0.2 0.0 le ng th re du ct io n l/l o 1.0 0.8 0.6 0.4 0.2 0.0 020406080100120 sample 1 20 10 0 50 2010 15 30 track length (µm) fi ss io ntr ac k ag e (m a) elapsed time (myr) final fission-track age = 22 ma80 0 20 40 60 95 806040200 95 fi ss io ntr ac k ag e (m a) n um be r o f t ra ck s n um be r o f t ra ck s 80 0 20 40 60 806040200 final fission-track age = 62 ma 020406080100120 k p e o m p 21 time (ma) time (ma) track length (µm) elapsed time (myr)time (ma) 40 0 20 60 80 100 120 140 160 200 180 te m pe ra tu re (° c ) tracks totally annealed 020406080100120 50 2010 15 20 10 0 30 a b c d fig. 20. thermal history response of fission tracks in apatite under geological conditions. a: notional thermal history for a sedimentary sequence that underwent progressive burial through the cretaceous to middle cenozoic, followed by cooling due to uplift and erosion commencing at 30 ma and completed by 20 ma, with minor reburial from 20 ma to the present day. the thermal histories of two samples are highlighted. b: track-length shortening vs. time for tracks produced at different times in samples 1 and 2. as temperature increases, the length of all tracks is progressively reduced, and because temperature dominates over time in the kinetics of annealing, at any time during this phase all but the most recently formed tracks at any given time have the same mean length (although each population of tracks has a finite spread in length). at the point when the maximum temperature is reached and the history changes from heating to cooling, all tracks formed up to that point in time are effectively ‘frozen’ at the length to which they have been reduced. they do not undergo further shortening because annealing rates are much slower at the reduced temperatures now prevailing, and they do not get longer because the annealing process is irreversible. those tracks formed after the onset of cooling remain longer because of the lower annealing rates at the prevailing lower temperatures. sample 2 reached a maximum temperature sufficient to reduce the length of all tracks produced up to that time to zero. at the present day, this sample contains only one track population, formed after the sample cooled to temperatures at which tracks could be retained (c. 110°c for typical apatite compositions). c: track-length distributions for samples 1 and 2 resulting from the thermal histories shown in a. for sample 1, two populations of tracks are present at the present day; a shorter population representing tracks formed up until the onset of cooling from the palaeo-thermal maximum, and a longer population formed after the onset of cooling. for sample 2, the measured track-length distribution will reflect the thermal history in the post-cooling period only. d: evolution of fission-track age with time resulting from the thermal histories shown in a. for sample 1, the final measured fission-track age will represent the summed contribution of the two populations of tracks; the shorter component will contribute a reduced component to the fission-track age compared to the time interval over which tracks have been retained. on the other hand, the contribution to the fission-track age of the longer population will be much closer to the time elapsed since the onset of cooling. for sample 2, the final fission-track age will be determined by the time when the sample began to retain tracks, but moderated by the degree of length reduction of tracks formed during the cooling history. while this is based on a mono-compositional apatite of durango composition using the laslett et al. (1987) model, the nature of this response is common to all forms of kinetic models, and is a fundamental property of the apatite fission-track system governed by a fanning arrhenius plot. myr: million years. modified from green et al. (2002). 39 4.1.4 extracting thermal history information from apatite fission-track data extracting explicit thermal history solutions directly from apatite fission-track data is not possible because of the high degree of redundancy in the data (i.e. many histories result in the same measured age and length parameters, fig. 24). instead, the problem is approached by forward modelling the apatite fission-track parameters expected from a range of specified thermal histories and defining the range of conditions which provide predictions that are consistent with the measured data. within this common philosophy, a range of approaches has been developed, each of which has its advantages and disadvantages. this disparity can often confuse the nonspecialist, so below we review the most commonly used approaches. most published studies represent a variation on one of these basic themes. the basics of data generation in all approaches are similar in most respects. geotrack approach (afta®): the approach adopted in our own studies, referred to as afta (apatite fissiontrack analysis), is designed primarily for application to sedimentary basins, and takes account of the fact that sedimentary horizons are deposited at the surface then stratigraphic age 0 1 2 3 d ep th (k m ) d ep th (k m ) a b c 0 1 2 3 palaeo-110°c isotherm present-day 110°c isotherm mean length (µm) 0 5 10 15 20 20 n um be r o f tr ac ks 10 0 30 track length (µm) fission-track age (ma) increasing degree of annealing of tracks formed prior to onset of cooling total annealing of tracks formed prior to onset of cooling total annealing of tracks in present-day thermal regime 0 20 40 60 80 100 120 8 9 10 11 12 13 14 15 20 n um be r o f tr ac ks 10 0 30 20 n um be r o f tr ac ks 10 0 30 20 n um be r o f tr ac ks 10 0 30 20 n um be r o f tr ac ks 10 0 30 fig. 21. predicted variation of fission-track parameters with depth for a vertical rock section with a thermal history of the style shown in fig. 20. a: fission-track age declines rapidly with increasing depth through the shallower section, as the mean length of the shorter population of tracks (formed up to the onset of cooling) is progressively shortened, and the proportion of these tracks that can reach the polished grain surface and be revealed, decreases. b: similarly, the mean track length reduces due to the decreasing mean length of the shorter population, evident in the track-length distributions (c). as the depth (and temperature of c. 110°c) corresponding to total annealing of all tracks formed prior to the onset of cooling is approached, the mean track length begins to increase again, as the shorter population of tracks becomes increasingly difficult to reveal and therefore contributes less to the overall mean for the sample, which is increasingly dominated by the longer population of tracks formed after cooling. as the transition from partial to total annealing of tracks formed prior to cooling is crossed at the palaeo-isotherm of c. 110°c, the mean track length increases abruptly as the sample is now dominated only by longer tracks formed after the onset of cooling. and the fission-track age reduction shows a characteristic ‘break-in-slope’, below which only a single component of tracks is present, with parameters controlled by the history after the onset of cooling. with further increase in depth, both the fission-track age and mean track length show progressive reduction to zero at a present-day 110°c, although in detail this temperature is controlled by apatite composition and the timescale of heating/burial. modified from green et al. (2002). 4040 stratigraphic age 300 200 100 0 fi ss io n tr ac k ag e (m a) 0 0.4 0.8 1.2 1.6 2.0 2.4 300 200 100 0 fi ss io ntr ac k ag e (m a) 0 0.4 0.8 1.2 1.6 2.0 2.4 300 200 100 0 fi ss io ntr ac k ag e (m a) 0 0.4 0.8 1.2 1.6 2.0 2.4 gc 440-11 95°c a b c de 300 200 100 0 fi ss io ntr ac k ag e (m a) 0 0.4 0.8 1.2 1.6 2.0 2.4 gc 440-14 107°c gc 440-17 121°c cl (wt%) cl (wt%) cl (wt%)cl (wt%) gc 440-18 124°c 14 17 18 fission-track age (ma) 0 50 100 0 1 2 3 4 20 40 60 80 100 120 140 flaxmans-1 d ep th (k m ) te m pe ra tu re (° c ) stratigraphic age stratigraphic age stratigraphic agestratigraphic age 11 fig. 22. a: variation of central fission-track age with depth for samples from the flaxmans-1 well, otway basin, se australia. b–e: the variation of fission-track age with chlorine content for individual apatite grains from four selected samples. in each of these samples the most sensitive (low-cl) grains are totally annealed (i.e. zero ft age) while more retentive (higher-cl) grains giving ages up to the depositional age and above. with increasing present-day, down-hole temperature (the maximum post-depositional temperature in these samples), the transition to total annealing shifts to progressively higher cl contents, demonstrating the systematic influence of chlorine content on annealing. note that while most of the central fission-track ages define a generally smooth decrease with increasing temperature, the age for sample gc440-14 is off trend. this is due to the absence of grains with cl >1.6 wt% compared to adjacent samples, which are dominated by more retentive grains and therefore give higher central ages. this major effect of chlorine on apatite fission-track age must be taken into account in order to extract meaningful geological constraints from apatite fission-track data. modified from green & duddy (2013). 41 buried and heated to some maximum depth and temperature, after which they may be exhumed and cooled. this allows a convenient parameterisation of the history in terms of one or more episodes of heating and cooling. because of the high level of redundancy in the data (see above), in order to extract useful information from afta data it is important to establish a framework with some fixed independent constraints, without which the range of possible solutions is unworkably large. for this purpose, we construct a ‘default thermal history’ (dth), which represents that part of the history that can be constrained from geological evidence. the dth for an outcrop sample of sedimentary rock is defined by two points, deposition and the present day, at which the sample is at the appropriate mean surface temperature. for basement outcrop samples, a similar approach can be adopted using the age of the oldest overlying sedimentary unit (if no cover is present, the intrusion age or youngest metamorphic age can be used). for downhole samples, the dth is calculated from the burial history derived from the preserved section combined with the presentday thermal gradient. if the afta data can be explained either solely by the dth or by a combination of the dth and tracks inherited from sediment source regions (for sedimentary rock samples), no further information can be obtained from the data because the data are dominated by the maximum temperature. but if the data show a higher degree of annealing (in terms of either fission-track age reduction or track-length reduction) than can be accommodated by the dth then the sample has been hotter during the period covered by the dth and the data can 0 2 4 6 8 10 12 14 16 18 m ea n tr ac k le ng th (µ m ) f(t,t) 1.0e+04 1.5e+04 2.0e+04 2.5e+04 3.0e+04 3.5e+04 4.0e+04 barbarand et al. (2003): all apatites except min bam: 4.67 wt% cl lin: 2.01 wt% cl gun: 1.54 wt% cl fct: 0.83 wt% cl ful: 0.71 wt% cl umb: 0.41 wt% cl dur: 0.39 wt% cl gil: 0.031 wt% cl far: 0.01 wt% cl drv: 0.01 wt% cl unk: 0.002 wt% cl wil: 0.00 wt% cl 0 2 4 6 8 10 12 14 16 18 a b m ea n tr ac k le ng th (µ m ) f(t,t) carlson et al. (1999): all apatites except hs, ay, pc, kp b3: 6.4 wt% cl b2: 2.95 wt% cl fc: 0.81 wt% cl durango: 0.43 wt% cl uk: 0.08 wt% cl wk: 0.04 wt% cl rn: 0.03 wt% cl ol: 0.03 wt% cl pq: 0.01 wt% cl sc: 0.01 wt% cl 1.0e+04 1.5e+04 2.0e+04 2.5e+04 3.0e+04 3.5e+04 4.0e+04 fig. 23. mean track lengths from laboratory annealing experiments reported by carlson et al. (1999) and barbarand et al. (2003), plotted against a unifying function of temperature (t) and time (t), which reduces all data to a common scale. this function is of the form f(t,t) = [log t – log to]/[(1/t)–(1/ to)], log to = –10 and 1/to = 0.001. apatites of different cl content are coded to illustrate this variation, with high cl contents (>1 wt% cl) shown in blue colours and large symbols, apatites low in cl (<0.1 wt % cl) shown in pale colours and smaller symbols, compositions around 0.4 wt% cl shown in yellow and those around 0.8 wt% cl in green colours. the durango apatite (c. 0.4 wt% cl, yellow) and fish canyon tuff apatite (c. 0.8 wt% cl, green circles) are common to both datasets. these results clearly illustrate the first order control on annealing rates exerted by cl content, with apatites high in chlorine giving longer lengths for any given heat treatment than those low in cl. while the first-order control from chlorine is clear, other elements produce additional variation, and several apatites from each dataset have been omitted as they are not consistent with the main body of data. however, in data from natural samples, no systematic compositional control on annealing other than chlorine has yet been identified. modified from green & duddy (2013). 4242 be used to define the main features of palaeothermal history (i.e. that period when the sample was hotter than it is today). in sedimentary rock samples which have not been heated above c. 60°c, the afta data may be dominated by tracks formed prior to deposition, and it may not be possible to resolve the effects of post-depositional heating. in such cases only an upper limit to the magnitude of the maximum post-depositional temperature can be obtained (green & duddy 2013). by modelling the expected fission-track age and tracklength distribution and their variation with wt% cl resulting from a range of possible thermal histories, we can define the range of values of maximum palaeotemperature and the onset of cooling resulting in predictions which match the measured data within 95% confidence limits (figs 25–27). the approach is based on likelihood theory similar to that described by gallagher (1995) but it also includes the influence of composition. a multicompositional annealing model is used which takes specific account of the influence of wt% cl. this model consists of a series of parallel equations, each taking the form of a linear fanning arrhenius plot with non-zero a 0 20 40 60 80 100 te m pe ra tu re (° c ) time (ma) 1.0 0.8 0.6 0.4 0.2 0.0 le ng th re du ct io n l/l o 5 10 15 200 track length (µm) 20 10 0 30 n um be r o f t ra ck s 100 80 60 40 20 0 time (ma) 100 80 60 40 20 0 b 0 20 40 60 80 100 te m pe ra tu re (° c ) time (ma) 100 80 60 40 20 0 c 0 20 40 60 80 100 te m pe ra tu re (° c ) time (ma) 100 80 60 40 20 0 1.0 0.8 0.6 0.4 0.2 0.0 le ng th re du ct io n l/l o time (ma) 100 80 60 40 20 0 1.0 0.8 0.6 0.4 0.2 0.0 le ng th re du ct io n l/l o time (ma) 100 80 60 40 20 0 5 10 15 200 track length (µm) 20 10 0 30 n um be r o f t ra ck s 5 10 15 200 track length (µm) 20 10 0 30 n um be r o f t ra ck s fig. 24. shortening trajectories (centre row) for tracks produced at different times through three thermal history scenarios (top row) representing increasing levels of complexity from a to c. the resulting track-length distributions are also shown (bottom row). despite the obvious differences in the thermal histories, the resulting track-length distributions are almost indistinguishable because both the time and magnitude of maximum temperatures and the rate of cooling from maximum palaeotemperatures are the same for each history (circled points on the thermal histories). this outcome reflects the fundamental kinetics of the apatite fission-track system such that the data are sensitive only to the magnitude of the maximum temperature and the timing of the onset of cooling (in relation to the overall time over which tracks have been retained), and preserve no information on the prior history (except that temperatures must have been lower than at the palaeo-thermal maximum). thus, successive heating episodes overprint the effects of earlier episodes, leaving only evidence of the maximum temperature episode and the subsequent history after cooling from the palaeo-thermal maximum. it is therefore not possible to discriminate between these three scenarios from apatite fission-track data. modified from green & duddy (2013). 43 fig. 25. principles of afta interpretation illustrated for a monocompositional apatite, showing how a thermal history solution can be extracted from measured afta parameters (fission-track age, mean track length and track-length distribution). for samples of sedimentary rock it is necessary to know the stratigraphic age and present temperature of the sample. by predicting the afta parameters for various thermal history scenarios, we can define the best-fit thermal history. as a first step, assume that cooling from the maximum palaeotemperature occurred at the midpoint of the history (120 ma in this case). a: by varying the maximum temperature and comparing measured and predicted parameters, we find a good match with the shorter population of tracks in the measured track-length distribution at a maximum palaeotemperature of 90°c. but the predicted track-length distribution contains too many long tracks. b: by fixing the maximum temperature at 90°c and varying the timing of cooling, a good match between the predicted and measured track-length distributions, as well as the fission-track age, is achieved with cooling commencing at 50 ma. note that no attempt is made to define the whole thermal history because the history prior to the onset of cooling is overprinted by the thermal maximum. note also that by itself, the measured fission-track age of 183 ± 12 ma provides no direct indication of the time of cooling, which only comes from kinetic modelling of the details of the track-length distribution together with the fission-track age. predictions based on a mono-compositional apatite of durango composition using the laslett et al. (1987) model. modified from green et al. (2002) and green & duddy (2013). measured track-length distribution cooling at 200 ma ft age = 213 ma mean length = 13.6 µm cooling at 10 ma ft age = 172 ma mean length = 10.7 µm n um be r o f t ra ck s max. temp = 70°c ft age = 219 ma mean length = 13.7 µm max. temp = 100°c ft age = 148 ma mean length = 12.4 µm max. temp = 90°c ft age = 196 ma mean length = 12.6 µm a. vary maximum palaeotemperature b. vary timing 20 0 40 60 80 100 120 140 basic data stratigraphic age: 240 ma present temperature: 10°c fission-track age: 183 ± 12 ma mean track length: 11.7 ± 0.2 µm te m pe ra tu re (° c ) time (ma) cooling at 50 ma ft age = 181 ma mean length = 11.5 µm 40 30 20 10 0 0 5 10 track length (µm) track length (µm) track length (µm) shorter peak too long shorter peak too short shorter peak about right: so fix the temperature and vary timing not enough short tracks not enough long tracks just about right! track length (µm) track length (µm) track length (µm) 15 20 n um be r o f t ra ck s 40 30 20 10 0 0 5 10 15 20 0 5 10 15 20 0 5 10 15 200 5 10 15 200 5 10 15 20200 100 0 20 0 40 60 80 100 120 140 te m pe ra tu re (° c ) time (ma) 200 100 0 n um be r o f t ra ck s 20 10 0 0 5 10 track length (µm) 15 20 4444 intercept, with constants which vary systematically with wt% cl. no attempt is made to define the whole thermal history because, as shown in figs 19, 20, 24, the history prior to the onset of cooling is overprinted by the thermal maximum. instead we focus on determining those aspects of the thermal history that directly control the measured afta parameters, viz. the maximum palaeotemperature and the time at which cooling from the palaeothermal maximum began (fig. 20). additional episodes of heating and cooling following the onset of cooling from the palaeothermal maximum can often be resolved, utilising the shortening of tracks formed after the initial cooling phase, provided that the magnitude and timing of the palaeothermal maximum and subsequent peak are sufficiently separated in temperature and time. in rare cases, three discrete episodes can be resolved in data from a single sample (e.g. turner et al. 2008). the principles involved are illustrated in fig. 27. while the episodic history of heating and cooling adopted in our approach is designed specifically for application to sedimentary basins, we believe that it is also relevant to basement terrains (green & duddy 2010; japsen et al. 2010; lidmar-bergström et al. 2013). such an approach is essential in basement regions where sedimentary outliers and re-exposed peneplains occur, revealing earlier cycles of exhumation, burial and re-exhumation (e.g. green & duddy 2006, 2007). examples of these situations are discussed in chapters 6 and 8. in routine application, afta data are commonly integrated with data from other palaeothermal indicators such as vitrinite reflectance (vr), and/or indicators of burial such as sonic velocity (e.g. japsen et al. 2007a). such data provide an independent check on the interpre0.0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0 100 200 300 400 a b thermal history scenarios: 80°c at 100 ma 100°c at 60 ma 120°c at 40 ma 0.0–0.1 wt% cl 0.1–0.2 wt% cl 0.2–0.3 wt% cl 0.3–0.4 wt% cl 0.5–0.6 wt% cl >0.6 wt% cl 0.4–0.5 wt% cl 0.0–0.1 wt% cl 0.1–0.2 wt% cl 0.2–0.3 wt% cl 0.3–0.4 wt% cl 0.4–0.5 wt% cl 0.5–0.6 wt% cl 0.6–0.7 wt% cl 0.7–0.8 wt% cl 40 30 20 10 0 0 5 10 track length (µm) 15 20 n um be r o f t ra ck s fi ss io ntr ac k ag e (m a) 0 5 10 track length (µm) cl (wt%) 15 20 0 5 10 track length (µm) 15 20 0 5 10 track length (µm) 15 20 0 5 10 track length (µm) 15 200 5 10 track length (µm) 15 200 5 10 track length (µm) 15 20 40 30 20 10 0 0 5 10 track length (µm) 15 20 n um be r o f t ra ck s measured track-length distribution in each compositional class track-length distribution in each compositional class predicted from the best-fit thermal history fig. 26. apatite fission-track interpretation methodology for a multi-compositional apatite sample. the same basic information and interpretation strategy as described in fig. 25 is used for samples which contain apatite of different compositions, but supplemented by the wt% chlorine of each apatite grain (measured by electron microprobe) in which a fission-track age or track length is measured. fission-track ages and track lengths are grouped according to the chlorine content into 0.1 wt% cl intervals, and a multi-compositional annealing model is used which takes specific account of the influence of wt% cl on annealing rates. the matching procedure is the same as for a sample of a single composition, but now involves the simultaneous matching of fission-track age and the details of the track-length distribution in all compositional groups present in the sample. in this example, 8 groups are present containing between 0.0 and 0.8 wt% cl. the best-fit match to the data is achieved for cooling from a maximum temperature of 100°c beginning at 60 ma (data in a permian sandstone from northeastern england). modified from green et al. (2002) and green & duddy (2013). 45 tation of the afta data, and ensure that the resulting thermal histories and the information on denudation derived from them are not affected by artefacts of any individual technique. final interpretations can also be integrated with information from sla (chapter 3), such as the presence of re-exposed peneplains and incised valleys. ‘hefty’ and related approaches: the majority of published apatite fission-track studies over the last 10 years or so rely on extracting thermal history information from the ‘hefty’ software package (ketcham 2005; ehlers et al. 2005) which is a development from the earlier ‘aftsolve’ software (ketcham et al. 2000). further details of analytical approach are provided by ketcham et al. (2007, 2009). in this approach, repeated forward modelling of data through various thermal history scenarios within specified limits results in definition of a range of viable histories for which predictions provide ‘acceptable’ fits to the measured data, including a more restricted range of solutions providing ‘good’ fits. the ranges of good and acceptable fits provide an assessment of the degree of uncertainty in the history. this approach typically allows for within-sample variation in annealing kinetics using etch pit size as the kinetic parameter, although wt% cl may be used (ketcham et al. 1999). the variation in kinetics with wt% cl embodied in this model is very similar to that in the geotrack model (above), and the two approaches give similar results when applied with similarly specified thermal history scenarios. either monotonic cooling or episodic heating and cooling can be employed, but in most published studies too much flexibility is built into the thermal history structure, resulting in wide ranges of allowed fits which reflect the redundancy issues discussed earlier (fig. 24), rather than providing meaningful constraints on the thermal history. this can be avoided to some extent by specifying thermal histories in a series of simple linear segments involving heating to a maximum temperature followed by cooling, although this is rarely done. temperature–time constraints can be imposed on the range of allowed fits, as boxes in temperature–time space through which the histories must pass. such constraints derived from geological evidence (e.g. from overlying sedimentary units) can narrow the range of likely histories, and in some ways can be thought of as analogous to the default thermal history approach outlined above. however, in practice details of such constraints, together with the underlying justification, are rarely provided. this is problematical, since these constraints can exert a critical control on the nature of the resulting thermal histories. the hefty approach provides solutions which are defined up to 200°c (with corresponding uncertainty limits), but this is purely an abstraction resulting from the assumed style of history, and solutions in this temperature range have no real meaning. note also that as this approach attempts to constrain the entire thermal history below 200°c, considerable computing time and effort is wasted because of the issue of redundancy highlighted in fig. 24. ‘monte-trax’ and related approaches: gallagher (1995) showed how forward modelling could be combined with a ‘genetic algorithm’ to constrain the range of thermal histories that is consistent with the measured data and to converge on the most likely history. this approach, embodied in the ‘montetrax’ software package, has commonly been used in basement terrains and crystalline rocks to constrain the complete thermal history below c. 110°c assuming monotonic cooling (although this is not essential). while this approach has been superseded by more sophisticated approaches (see below), it remains of interest because it underlies several key studies of denudation histories of the southern africa and brazil continental margins discussed in chapter 6, which are still the subject of considerable discussion. the kinetic model of laslett et al. (1987) is commonly employed in this approach, on the assumption that the durango apatite (young et al. 1969), for which this model was developed, is representative of common apatites. in many studies using this approach, apatite fission-track data have been used in isolation as the only constraint on thermal and denudation histories. adoption of the laslett et al. (1987) model as the keystone of this approach introduces some problems, particularly with the low temperature (<60°c) aspect of thermal history solutions, which commonly show evidence of major cooling within the last 10 to 20 million years. this is commonly regarded as spurious, and attributed to inaccuracy in predictions of the model at low temperatures (e.g. hendriks & andriessen 2002). however, a large contribution to this effect may arise from the fact that the durango apatite is not representative of apatite grains commonly analysed from basement rocks, which typically contain <0.1 wt% cl. in contrast, the durango apatite contains c. 0.4 wt% cl, and is thus more resistant to annealing than typical low-cl basement-derived apatites (ketcham et al. 1999; barbarand et al. 2003), which therefore show a greater degree of annealing compared to the predictions of a model based on durango apatite. green (2004) suggested that this can readily explain the ‘anomalous 4646 miocene cooling’ commonly seen in ‘monte-trax’ solutions. unfortunately, despite the advent of kinetic models which explicitly include the influence of chlorine content on annealing rates (e.g. ketcham et al. 1999), these have not been be incorporated into the monte-trax software. in an attempt to account for these problems, gunnell et al. (2003) renormalised the laslett et al. (1987) model. but in addition to modifying the low temperature behaviour, this also made the model more sensitive at high 400 300 200 100 0 0.80.0 0.2 0.4 0.6 1.0 1.2 1.4 1.6 1.8 cl (wt%) stratigraphic age fi ss io ntr ac k ag e (m a) default thermal history and predictions best-fit thermal history and predictions default thermal history and predictions best-fit thermal history and predictions prediction without second event prediction without earliest event track-length distribution predicted without the most recent event te m pe ra tu re (° c ) n um be r o f t ra ck s 0 150 200 100 50 maximum palaeotemp. from vr data outlier: excluded from detailed analysis tracks formed prior to the first palaeothermal maximum are totally annealed tracks formed between the palaeo-thermal maximum and the onset of cooling from the second palaeo-thermal peak contribute the shorter tracks in the length distribution tracks formed after the second cooling episode from the third palaeo-thermal peak constitute the longest tracks in the mode of the length distribution 1.0 0.9 0.8 0.7 0.6 0.5 0.4 0.3 0.2 0.1 0.0 le ng th re du ct io n b c d a (n.b. track paths illustrate behaviour in only a single apatite composition) measured track-length distribution track length (µm) 200 20 10 0 30 40 0280350 210 140 70 time (ma) 0280350 210 140 70 time (ma) 5 10 15 track length (µm) 200 20 10 0 30 40 5 10 15 fig. 27. definition of three events from afta data in one sample. this example, based on afta data from the dodo canyon k-03 well in the mackenzie valley, northwest territories, canada (after green & duddy 2013), illustrates how multiple events can be resolved from afta data, where the events are sufficiently separated in temperature and time. a: the variation of fission-track age with wt% cl over a wide range of chlorine contents in this sample provides definition of the two earlier events (colour scale as in fig. 26), while the most recent event is defined from track-length data (c and d). b: three palaeo-thermal episodes are required to explain all aspects of the afta data in this sample. the best-fit thermal history results in predictions which provide an excellent fit to both the variation of fission-track age with wt% cl (a) and also the measured track-length distribution (c). omitting either of the two earlier events results in a failure to fit the trend of age vs. wt% cl (a), while omitting the most recent event produces a poor fit to the longer end of the track-length distribution (c). d: predicted track-shortening trajectories for the best-fit history (b) illustrate the development of afta parameters through this history. all tracks are erased in the earliest episode, which leads to resetting of fission-track ages across the range of wt% cl. a population of shorter tracks is then produced at the second palaeo-thermal peak (note that the sedimentary section deposited after the initial cooling shows that re-heating is required). finally a third palaeo-thermal episode produces another population of tracks shortened to a lesser degree (constituting the longer part of the main mode of the distribution). a few longer tracks are produced after this final episode, but contribute only a small proportion of tracks. vr (vitrinite reflectance) data provide independent verification of the magnitude of the earliest episode. modified from green & duddy (2013). 47 temperatures, and results obtained using this modified model are therefore unreliable. a number of variations of this approach have been developed (e.g. corrigan 1991; lutz and omar 1991; willett 1997), some of which implement the laslett et al. (1987) annealing model, while others have used alternative models (e.g. crowley et al. 1991). however, extrapolation of these alternative models to geological timescales is not as accurate as that of the laslett et al. (1987) model, and these models have not been widely adopted. more recently gallagher (2012) described a new approach for inverse modelling of thermal history from apatite fission-track data using a bayesian transdimensional markov chain monte carlo approach, allowing the use of multiple samples and different styles of thermal history scenario, as well as combination of data from multiple techniques. cogné et al. (2012) demonstrated the use of this approach in samples from the brazil continental margin. fission-track stratigraphy: the qualitative use of apatite fission-track age patterns with depth or elevation to identify offsets between crustal blocks was pioneered in the transantarctic mountains by fitzgerald & gleadow (1987). this approach has been adopted more recently in other regions, e.g. the pyrenees (fitzgerald et al. 1999) and scandinavia (redfield et al. 2005). while this approach is largely free of any problems associated with the detailed thermal response, it fails to make full use of the data to provide quantitative thermal history constraints, and is susceptible to the influence of changes in apatite composition through the section (lorencak et al. 2004). as with most of the other approaches described above, apatite fission-track data are usually used in isolation in this approach. 4.1.5 monotonic cooling vs. episodic heating and cooling as discussed above, thermochronology studies of basement terrains are commonly carried out within a framework involving monotonic cooling from above c. 110°c to surface temperatures. in such studies, sedimentary outliers are often ignored (or avoided), and relief differentiation as evidence of former cover (now removed) in basement terrains is generally not discussed. the presence of the merest veneer of sedimentary cover lying on basement provides the key constraint that the underlying basement was at the surface when that cover was deposited. several studies have failed to take account of such basic geological constraints. persano et al. (2006) reported cooling histories for outcrop samples from the carboniferous bathurst batholith (eastern australia) involving protracted cooling from c. 110°c in permian times, but failed to allow for the presence of upper permian sedimentary outliers overlying the outcropping granite. these sedimentary remnants show that the granite was rapidly exhumed following intrusion and was then reburied, which completely changes the interpretation of the data and the resulting conclusions regarding the development of the landscape. despite this error being pointed out by brown (2007), gibson (2007) and green & duddy (2007), in their response persano et al. (2007) did not acknowledge the significance of these sedimentary outliers. similar situations where geological constraints have not been taken into account, leading to erroneous interpretations, include johnson & gallagher (2000; see japsen et al. 2010), persano et al. (2005), and pedersen et al. (2012; see japsen et al. 2013b). these and other examples where geological constraints have important implications for the interpretation of apatite fissiontrack data are discussed further in section 8.2. while episodic heating and cooling is clearly the most realistic scenario for sedimentary basins, as witnessed by the common occurrence of unconformities in sedimentary sequences, such histories are also reasonable for basement terrains where sedimentary outliers are present. similar comments apply where the present-day relief reveals old surfaces which have been preserved beneath a former cover (section 3.1.2). since episodic histories apply in these basement regions, there seems to be no reason why this should not also be true in basement areas devoid of present-day cover, which could simply indicate that the former cover has been totally stripped. this is particularly the case along epcms where stratigraphic landform analysis suggests the presence of re-exposed peneplains. recognition of the presence of sedimentary outliers can lead not only to major changes in interpreted cooling rates and corresponding denudation rates, but also to more fundamental aspects of geological evolution. weber et al. (2005), in a study of outcropping basement samples from the yilgarn craton of western australia, showed that the preservation of thin permian sedimentary remnants required rapid late carboniferous to early permian exhumation of basement to the surface and subsequent kilometre-scale reburial prior to renewed late palaeozoic – mesozoic exhumation. this study therefore revealed 4848 the presence of a thick former sedimentary cover over this supposedly stable cratonic region. this emphasises the common applicability of histories involving episodic cooling and reheating, even in basement regions. the south swedish dome (section 3.1.2) is another key area where combining thermochronology with constraints from geology and sla defines a history of episodic heating and cooling (burial and exhumation) through the phanerozoic. cambrian strata lie directly on basement in the north and east while jurassic and mainly cretaceous strata rest on basement in the south and west (lidmar-bergström 1988, figs 1, 2; japsen et al. 2002, fig 15). published apatite fission-track data (cederbom et al.2000; cederbom 2001) define late palaeozoic cooling from temperatures around 100°c or above, reflecting erosional removal of a substantial thickness of palaeozoic rocks deposited following formation of the sub-cambrian peneplain and deposition of remnant cambrian sediments. the apatite fission-track results also require oligocene–miocene cooling, following deposition of preserved mesozoic units, implying deposition of upper cretaceous to palaeogene cover. as illustrated by japsen et al. (2002) this cover preserved the hilly relief surface produced during mesozoic weathering. further support for the view that episodic heating and cooling provides a preferred framework for interpreting low temperature thermochronology data compared to monotonic cooling comes from a study by flowers & kelly (2011) of basement cores from a borehole through thin sedimentary cover overlying precambrian basement of the us mid-continent region. based on a combination of apatite fission-track data and apatite (u-th)/he dating they showed that basement samples showed evidence of major heating and cooling within each of several major regional unconformities (cambrian–permian, permian–cretaceous and cretaceous–recent), defining a history of episodic heating and cooling similar to those described above. although only c. 200 m of sedimentary cover represents phanerozoic time at the borehole location in kansas, the presence of this thin cover demonstrates that the underlying basement was not continuously exhumed through phanerozoic time, and monotonic cooling is therefore inappropriate. flowers & kelly (2011) used the presence of the cover units to define periods at which the basement rocks were close to the surface. only a slightly greater degree of erosion over the last 100 million years or so would have totally removed all the evidence that confirms this episodic history, though the real history would remain the same. this needs to be borne firmly in mind in considering likely histories for basement terrain devoid of sedimentary cover. 4.1.6 the meaning of a fission-track age in early studies, fission-track ages from crustal sections with simple histories were commonly discussed in terms of a zonation of ages (e.g. naeser et al. 1989). at shallow depths, ages were regarded as unaffected (‘zone of no annealing’, <70°c), while at depths greater than c. 3–4 km (temperatures in excess of c. 125°c), no tracks are retained (‘total annealing zone’). between these two extremes, fission-track ages are progressively reduced to zero through a ‘partial annealing zone’ (paz). while this zonation provided a simple conceptual basis for early studies, the combination of borehole data and laboratory experiments discussed above showed that reduction in both fission-track age and track length proceed at temperatures below 70°c, albeit more slowly than at higher temperatures. this can be seen most easily in confined track-length data due to the higher precision of these measurements compared to fission-track ages (fig. 18). the true meaning of a fission-track age can only be assessed in conjunction with the distribution of confined track lengths (e.g. gleadow et al. 1986a, b). an apatite fission-track age will only provide a direct indication of the time over which tracks have been retained in a sample which has cooled extremely rapidly from >110°c to temperatures less than 50°c and subsequently remained at such temperatures. in such samples, track lengths will be similar to those in the age-standard apatites used to calibrate the method (section 4.1.1). experience has shown that such situations are relatively rare. apatite fission-track ages therefore only rarely indicate the time over which tracks have been retained, and must be assessed in tandem with track-length data before the true significance becomes clear. in the same way, a fission-track age rarely reflects a ‘cooling age’ (in the sense that cooling occurred at a particular time), and we recommend that this term should not be used. in slowly cooled terrains, ages are often referenced to a ‘closure temperature’, below which the daughter product is effectively retained. however, literal interpretation of this concept is of dubious validity for apatite fission-track ages because tracks formed throughout the cooling history undergo significant shortening (fig. 19) and their contribution to the fission-track age 49 will consequently be reduced compared to the overall time over which tracks have been retained. in general, an apatite fission-track age should be regarded simply as a thermal history parameter, representing the integrated balance between the production of tracks by spontaneous fission and the reduction in the probability of tracks intersecting a polished and etched grain surface due to the reduction in track length resulting from the thermal history (figs 20, 21). 4.1.7 ‘boomerang plots’ the meaning of an apatite fission-track age is further illustrated by results from north-west england. green (1986) reported results in samples from palaeozoic intrusions, many of which were unroofed by the end of the devonian, and which have subsequently undergone essentially a single dominant episode of heating and subsequent cooling. the results define a systematic relationship between mean confined track length and fission-track age (fig. 28), defining a boomerang-like trend. samples that have undergone little thermal disturbance since the palaeozoic have old ages with relatively long 0 100 200 300 15 14 13 12 11 10 track length (µm) a b c d n um be r o f tr ac ks fully reset ages partly reset ages original (unreset) ages fission-track age (ma) m ea n tr ac k le ng th (µ m ) track length (µm) track length (µm) track length (µm) track length (µm) track length (µm) track length (µm) n um be r o f tr ac ks track length (µm) track length (µm) track length (µm) track length (µm) track length (µm) 0 100 200 300 15 14 13 12 11 10 fission-track age (ma) m ea n tr ac k le ng th (µ m ) fig. 28. boomerang plot: relationship between mean track length and fission track age for a suite of outcrop samples from north-west england which have undergone cooling from different maximum palaeotemperatures at the same time (modified from green 1986). a: measured track length distributions. b: measured data. c: illustration of track-length distributions as mixtures of two populations. d: illustration of the evolution of afta parameters with increasing maximum temperature (from right to left). samples that experienced low maximum temperatures have old ages with relatively long mean track lengths, while samples in which all fission tracks were totally annealed prior to the onset of cooling give much younger fission-track ages (‘reset ages’), also characterised by long (c. 14 µm) mean lengths. between these two extremes, as the fission-track age decreases (representing increasing maximum palaeotemperatures prior to the onset of cooling) the mean track length decreases as the partially annealed tracks are progressively shortened. this continues until the final stages of age reduction, when the partially annealed tracks become so short that their contribution to the mean length is diminished and the mean length increases with further reduction in fission track age, trending upwards towards the long mean length characterising the reset ages. this dataset can be considered analogous to the situation illustrated in fig. 21, but with all samples cooling to low (near surface) temperatures, such that each contains a population of long tracks formed after cooling.modified from green et al. (2002) and green & duddy (2013). 5050 mean track lengths, while samples in which all fission tracks were totally annealed prior to the onset of cooling gave much younger fission-track ages (‘reset ages’), also characterised by long (c. 14 µm) mean lengths. between these two extremes, as the fission-track age decreases (representing increasing maximum palaeotemperatures prior to the onset of cooling), the mean track length decreases as the partially annealed tracks are progressively shortened. this continues until the final stages of age reduction, when the partially annealed tracks become so short that their contribution to the mean length is diminished and the mean length increases with further reduction in fission-track age, trending upwards towards the long mean length characterising the reset ages. this trend can be viewed as analogous to the depth trend in fig. 21, except that all samples have a common long component because they all cooled to near-surface temperatures in the case of a suite of outcrop samples. in contrast, age vs. length data from many other regions show a very different type of trend to the simple pattern reported by green (1986). for example, results from norway (rohrman et al. 1995) show an almost opposite relationship to the classic ‘boomerang’ trend, while data from africa and brazil (gallagher & brown 1999b) show wide dispersion with only a slight tendency towards longest lengths associated with the youngest ages. compared to the simple situation in north-west england where a single dominant heating/cooling episode has produced a well-defined trend, such relationships imply a much more complex history, most likely involving a series of cooling episodes, each of which may vary in magnitude across the region. therefore, such plots should be interpreted with care, and detailed attention should be paid to the systematic change in the form of the track-length distribution through the plot, which was central to the original description by green (1986), but which is often overlooked in other studies. gallagher & brown (1997) reiterated the usefulness of plotting mean track length vs. fission-track age in considering the implications of regional apatite fission-track datasets, but within the context of monotonic cooling histories. this can lead to misconceptions, as explained in section 4.1.9. 4.1.8 uplift rates from apatite fission-track age profiles in early studies, the increase of fission-track age with elevation in mountainous regions was interpreted in terms of monotonic cooling, representing the progressive closure of the system and retention of fission tracks as samples moved upwards in the rock column and cooled below the closure temperature. such age patterns were used to determine rates of uplift and erosion (e.g. wagner & reimer 1972) based on the assumption that the change in apatite fission-track age with elevation represented slow cooling associated with progressive denudation of the mountain range. on this basis, segments of the age vs. elevation trend with different slopes were interpreted as representing a change in the rate of denudation. however, subsequent integration of confined track-length information with the fission-track age data (fitzgerald & gleadow 1987) showed that the upper segment of such trends above the breakin-slope in fact represents an uplifted (exhumed) partial annealing zone (cf. fig. 21), while samples below the break-in-slope were totally annealed prior to the onset of exhumation. thus, the slope of the age vs. elevation profile cannot be interpreted simply in terms of slow cooling. instead, the decrease with age through the upper section towards the break-in-slope represents increasing degrees of annealing of deeper samples prior to the onset of cooling. 4.1.9 long term residence in the partial annealing zone vs. heating and cooling it is often asserted that the presence of an exhumed partial annealing zone, as identified by a break-in-slope in the variation of fission-track age with elevation or depth (fig. 21), represents a prolonged period of residence in the partial annealing zone prior to exhumation. this again reflects an attitude framed in terms of monotonic cooling. consideration of fig. 21 shows that long-term residence is not necessary, as data of this type are easily produced by heating (e.g. by burial) of the sequence to temperatures characterising severe annealing followed immediately by rapid subsequent cooling/exhumation. this misunderstanding reflects the viewpoint of many workers that monotonic cooling histories form the most appropriate framework for interpreting apatite fissiontrack data. in contrast we submit that episodic heating 51 and cooling histories are more relevant in many situations (cf. japsen et al. 2010; green & duddy 2010). 4.1.10 limitations of apatite fission-track methods as noted above, failure to recognise the limitations of the apatite fission-track system is a common cause of errors in interpreting the data. one major limitation is the fact that the data are sensitive only to key moments in the overall history. for episodic heating and cooling histories, such as illustrated in figs 20 and 21, it should be clear that the data contain no information on the variation of temperature during the heating phase prior to the onset of cooling. it is for this reason that application of afta is focused specifically on determining the magnitude of the maximum temperature and the time at which cooling begins (section 4.1.4). in addition, factors such as the inherent spread in the width of the track-length distribution, resulting from the energetics of the fissionprocess (green 1980), and the number of tracks counted in determining a fission-track age impose fundamental limits to the temperature and time resolution that can be achieved. a resolution of 10°c in temperature is possible for the 60 to 90°c range (equivalent in broad terms to a resolvable difference of around 1 µm in mean track length over this temperature range) while 5°c may be achievable between 90 and 110°c, due to increasingly rapid decrease in fission-track age. for samples that have been heated to less than 60°c, it is often possible only to set an upper limit to the magnitude of heating, particularly in sedimentary rocks which contain tracks formed prior to deposition. in samples that have been heated to temperatures above c. 110°c where tracks are totally annealed, the history prior to cooling below this temperature can of course not be constrained. poissonian uncertainties in fission-track age (typically around 5% at best) derived from the numbers of tracks counted set a fundamental limit to resolution in time. the consequence of these considerations is that for histories involving a series of heating and cooling episodes, afta will reveal only the major episodes, and may give solutions that represent the unresolved effects of multiple episodes. solutions framed in terms of monotonic cooling, in situations where the real histories involved a series of episodes of heating and cooling, will represent only a broad approximation to the dominant palaeothermal episodes. given these factors, combined with the redundancy in the data (fig. 24), it should be appreciated that in all approaches to extracting quantitative thermal history constraints from apatite fission-track data, the full detail of the underlying thermal history can never be defined. the resulting thermal history solution from any approach should therefore be considered as providing an approximation to the true history and the skill in using these techniques is to recognise the range of possible solutions that can be accommodated by a given set of data. in this context, it should be appreciated that ranges of uncertainty attributed to cooling histories derived from other approaches described in section 4.1.4 are contingent on the assumed form of the history (usually monotonic cooling), and alternative styles of history may be possible (e.g. episodic heating and cooling) for which other uncertainty limits will apply. 4.1.11 alternative views of fission-track annealing while the role of thermal annealing in controlling the evolution of apatite fission-track age and length is widely acknowledged, some workers have suggested that other processes can also influence data in some circumstances. wendt et al. (2002) suggested that contrary to previous ideas (e.g. fleischer et al. 1975) pressure could play an important role in affecting the kinetics of the annealing process. kohn et al. (2003) criticised many aspects of the ideas presented by wendt et al. (2002), and despite contrary arguments by vidal et al. (2003), the idea that pressure can affect fission-track annealing kinetics is not a widely held view. we note that if pressure were to have a significant effect on annealing rates, it might be expected to be most obvious in results from deep drill holes in cratonic regions characterised by low thermal gradients, but results to date from such regions (e.g. rohrman 1995; lorencak et al. 2004) show no such evidence. more recently, hendriks & redfield (2005) suggested that apatite fission-track ages in finland were anomalously young because of a non-thermal annealing process due to the effects of a long-term radiation dose from alpha decay over millions of years. this suggestion, if true, has severe implications for the interpretation of apatite fission-track data. one of the lines of evidence used to justify this claim was the observation that apatite fission-track ages were younger than (u-th)/he ages in the same apatites, contrary to expectation (see next sec5252 tion). but green et al. (2006) showed that this is due to anomalous behaviour in the (u-th)/he system, and not to anomalous fission-track annealing behaviour. other evidence used to support the notion of ‘radiationenhanced’ fission-track annealing can be interpreted in conventional fashion in terms of standard processes of thermal annealing (green & duddy 2006; larson et al. 2006) and the concept of radiation-enhanced annealing currently finds little support. 4.1.12 summary the thermal response of fission-tracks in apatite is well understood (at least empirically), and provides results which are consistent with constraints from other techniques when applied in tandem (section 4.3). despite often relatively low precision on thermal history constraints, the technique provides unique information yielding interpretations that cannot be obtained from other methods. while approaches used by different groups may differ in detail, the response of the system is well defined, and thermal history information provided by the technique should be reliable within the constraints and limitations of the methods used. perceived problems with the technique, as discussed for example by gunnell (2000), are likely to result from a failure to appreciate these limitations, rather than any fundamental problem with the technique itself. in many cases too much is expected from the technique. when attention is focused on the unique information that can be obtained, and results are integrated with regional geological constraints, the power of the technique becomes apparent. 4.2 apatite (u-th)/he dating 4.2.1 historical background the development of apatite (u-th)/he dating (zeitler et al. 1987; wolf et al. 1996) potentially opened up a lower temperature window for thermochronology. whereas fission-track dating is based on spontaneous fission-of 238u atoms, (u-th)/he dating is based on the alpha decay of uranium and thorium isotopes (and to some extent samarium, e.g. grist & zentilli 2005). alpha particles, being helium nuclei, are neutralised rapidly after emission from the parent nucleus, and helium gas accumulates in the apatite lattice at a rate dependent on the u and th content of the apatite. but once formed, the he gas is progressively lost from the apatite due to diffusion at a rate which depends on temperature, in similar fashion to the annealing of fission-tracks. this balance between the production and loss of he forms the basis of this technique. initial studies (e.g. house et al. 1997; warnock et al. 1997; wolf et al. 1997) illustrated the potential of the technique to provide useful information at temperatures in the 40–80°c range, possibly giving this method even greater potential than apatite fission-track methods for providing information relevant to the interpretation of landscape development, and suggesting that the combination of the two techniques should be particularly powerful. 4.2.2 early success early applications of the method to elucidate thermal history information in geological conditions were based on kinetics of he diffusion in apatite and the way that this depends on temperature and grain size derived from laboratory experiments on durango apatite (farley 2000). in similar fashion to apatite fission-track methods, comparison of measured ages with values predicted from various thermal history scenarios allows definition of the range of thermal histories giving predictions that are consistent with the results. comparison of (u-th)/ he ages with apatite fission-track data from boreholes in the otway basin, south-east australia (house et al. 1999, 2002) illustrated the unique sensitivity of (u-th)/ he ages in apatite to temperatures in the range 20 to 80°c, and appeared to confirm the extrapolation of laboratory studies. integration of apatite (u-th)/he ages with afta data from the fresne-1 well in the taranaki basin, new zealand (crowhurst et al. 2002) provided further demonstration of the consistency between the two techniques, and illustrated the potential power of the combination of the two methods when applied to vertical sequences of samples, where the variation of (u-th)/he age with depth helps to further restrict the range of viable solutions defined from afta (fig. 29). apatite (u-th)/he dating has now been applied with apparent success in a wide range of environments from mountain belts to sedimentary basins. but while it was initially assumed that the he diffusion systematics from farley (2000) could be applied to all common apatites, it 53 has become clear more recently that this is not the case, as explained below. 4.2.3 evidence of greater complexity despite the early success, later studies (e.g. fitzgerald et al. 2006; hansen & reiners 2006; danisik et al. 2008) showed increasing evidence of excess dispersion in apatite (u-th)/he ages over what could be explained in terms of known controls (principally grain size). in addition to this random dispersion, studies in which (u-th)/ he data were integrated with afta (green & duddy 2006; green et al. 2006) revealed a systematic discrepancy between the two techniques which increased as either the (u-th)/he or the fission-track age increases (fig. 30). green et al. (2006) interpreted this trend to 25 µm 50 µm 75 µm 100 µm grain radii he age (ma) d ep th (k m ) 2.5 2.0 1.5 1.0 0.5 0 ft age and he age (ma) 0.5 1.0 0 1.5 2.0 2.5 time (ma) te m pe ra tu re (° c ) 01020304050607080 time (ma) 01020304050607080 time (ma) 01020304050607080 time (ma) 01020304050607080 0 50 100 150 0 50 100 150 0 50 100 150 0 50 100 150 d ep th (k m ) linear heating and cooling ba two cooling eventsprotracted coolingrapid cooling f ihg 00 10 20 30 40 50 60 70 80 90 2 4 6 8 10 he age (ma) 2.5 2.0 1.5 1.0 0.5 0 d 0 2 4 6 8 10 2.5 2.0 1.5 1.0 0.5 0 e 0 2 4 6 8 10 d ep th (k m ) 2.5 2.0 1.5 1.0 0.5 0 c 0 2 4 6 8 10 (u-th)/he age fission-track age stratigraphic age fig. 29. apatite (u-th)/he and fission-track ages in samples from the fresne-1 well, taranaki basin (new zealand). a: measured (u-th)/he ages are much younger than fission-track ages in the same samples. b, c, d, e: modelling the variation of (u-th)/he age with depth using four different thermal history scenarios (f, g, h, i), based on the interpretation of afta data from these samples, allows refinement of the preferred thermal history reconstruction, favouring a scenario involving two-stage inversion (e, i). ft: fission-track. modified from crowhurst et al. (2002). 5454 reflect enhanced he retentivity due to increasing levels of radiation damage in the apatite lattice, which results in anomalously old (u-th)/he ages compared to the expected response based on the farley (2000) systematics in samples with ages older than c. 50 ma. shuster et al. (2006) reported laboratory diffusion studies which confirmed the increase in helium retentivity as the amount of radiation damage in the apatite lattice increases. subsequent developments by shuster & farley (2009) and flowers et al. (2009) have led to a more advanced understanding, in which modelling of he diffusion is integrated with modelling of fission-track retention as a surrogate for the degree of radiation damage (the ‘rdaam’ model). in this model, the fission-track density is adopted as a proxy for the degree of radiation damage, and elimination of this damage is governed by fission-track annealing kinetics. this suggests that the apatite (u-th)/he method is best employed in tandem with apatite fission-track data. gautheron et al. (2009) provided further documentation of the influence of radiation damage in increasing he retentivity in apatite, and also offered an alternative quantitative model. these approaches offer considerable promise for providing more consistent thermal history solutions from multiple methods, but as illustrated by flowers & kelly (2011), detailed investigation is required to ensure that the data conform to the behaviour expected by the model that is used to extract thermal history information from the data. flowers & kelly (2011) provided guidelines for the interpretation of ‘over-dispersed’ apatite (u-th)/he data, including evaluation of the relationships between (u-th)/he age and equivalent uranium content and grain size. where such relationships are clear they can be used to control viable thermal history solutions. where no such relationships exist, flowers & kelly (2011) recommended that such data should not be used for thermal history interpretation and further data should be acquired to explain the dispersion of the results. for datasets where such detailed investigation has not been undertaken, we suggest that apatite (u-th)/he results should be treated with extreme caution, particularly when no apatite fission-track data are available from the same samples. the results presented by green et al. (2006) show that (u-th)/he ages in samples where the product of the age (in ma) and the equivalent uranium content in ppm is greater than c. 2500 cannot be explained in terms of the farley (2000) systematics (fig. 30), which should only be applied to young, low-uranium apatites. note that this applies even though these older ages often show a high degree of reproducibility. in the absence of quantitative modelling, (u-th)/he ages could, in principle, be used in a qualitative sense, similar to the ‘fission-track stratigraphy’ approach to apatite fission-track studies described above, by comparing age with elevation trends. but in addition to the systematic effects described by green et al. (2006), the excessive scatter in he ages, beyond that expected from purely analytical uncertainties described earlier, can cause serious problems in comparing he ages from different samples. [ u (ppm) + 0.24 × th (ppm) ] × ft age m ea su re d (u -t h) /h e ag e / p re di ct ed (u -t h) /h e ag e w greenland 861-2 lake mtn snowdon mud tank yilgarn rd8-39 durango fresne-1 0.0 0.5 1.0 1.5 2.0 2.5 3.0 3.5 4.0 4.5 4.0 10 0001000100 100 000 fig. 30. systematic change in he retention characteristics of apatite with increasing radiationdamage levels. the ratio of measured apatite (u-th)/he age to the value predicted from the best-fit t(t) solution derived from afta for each grain from a number of samples with well-defined thermal histories is plotted against the product of the measured fission-track age and a function of uranium and thorium contents of each grain (after green et al. 2006). this term provides an approximate measure of the number of alpha particle decays from u and th over a time equal to the measured fission-track age. these results show a progressive departure from consistency, highlighting the enhanced he retentivity of apatite as the degree of radiation damage in the crystal lattice increases (green et al. 2006). ft: fission-track. 55 even in the absence of such effects, because of the systematic effect of increasing levels of radiation damage, it is essential that only apatites with similar u contents and grain size should be compared. 4.2.4 he-closure temperatures and ‘cooling ages’ in some studies, (u-th)/he ages have been interpreted as representing the time at which samples cool through a closure temperature, typically around 50°c although in detail this will vary with radiation damage (shuster et al. 2006). but as highlighted by green & duddy (2006), open system behaviour is to be expected in the upper 2 km of the crust, and long-term residence at low temperatures, or even histories involving re-burial followed by exhumation, can produce results that can be easily mistaken for slow cooling. in practice, interpretations of apatite (u-th)/he ages are most reliable when accompanied by geological constraints (e.g. sedimentary outliers) and apatite fission-track data interpretations. 4.3 converting thermal history information to denudation history 4.3.1 denudation histories in principle, conversion of thermal histories to denudation or exhumation histories is simply a matter of dividing the thermal history (actually the difference between the sample temperature and the appropriate surface temperature) by the appropriate value of thermal gradient throughout the history. this provides the variation through time of the depth of the sample with respect to the surface. in a continuous cooling scenario, this is referred to as the denudation or exhumation history. many studies explicitly set out to determine long-term denudation histories of basement regions over hundreds of millions of years, within a framework of continuous cooling. in practice, this process is not so simple, because as noted in section 4.1 the thermal history is only controlled at certain key points in the history (fig. 24). outside these times, the thermal history is effectively unconstrained and a wide range of alternative temperatures are allowed by the data. for this reason, rigorous control on thermal gradients is usually possible only at limited times in the history. even if a constant heat flow is assumed, because the thermal history is effectively constrained only at key times the same is true of the resulting denudation history. long-term denudation rates derived from such studies therefore have little meaning. 4.3.2 palaeogeothermal gradients and removed section palaeotemperatures in a specific palaeothermal episode derived from afta data in samples over a range of depths (in sub-surface samples) or elevations (in outcrop samples from vertical rock sections) can be used to provide direct constraints on the palaeogeothermal gradient at the palaeothermal maximum. extrapolation of the fitted gradients to an appropriate palaeosurface temperature then provides an indication of the amount of additional section that was present at the palaeothermal maximum (fig. 31a). bray et al. (1992) described how statistical procedures can be used to define the range of palaeogeothermal gradients and amounts of missing section that are consistent with palaeotemperature constraints within 95% confidence limits. the inverse correlation between these two parameters results in a hyperbolic ellipsoid region of allowed values (fig. 31b). the accuracy of estimates of removed section that can be derived using this approach, within the available uncertainty limits, has been confirmed in a number of different situations where independent estimates are available of the amount of former cover that has been removed (e.g. green et al. 1995; crowhurst et al. 2002; japsen et al. 2007a). it should be stressed that estimating amounts of removed section by extrapolating a linear palaeotemperature profile assumes that the additional section had the same average thermal conductivity as the preserved section. if independent evidence suggests that this assumption is not appropriate, then a more detailed analysis using suitable thermal conductivities is required in order to provide a more accurate solution (fig. 32). but experience suggests that in most cases sedimentary sections are so heterogeneous that any influence of varying thermal conductivity is ‘smoothed out’ and a linear approximation for the palaeogeothermal gradient profile is acceptable. it should also be emphasised that the approach illustrated in fig. 31 only provides a constraint on the palaeogeothermal gradient at the palaeothermal maximum, 5656 and variation through time is largely unconstrained, unless multiple palaeothermal episodes can be recognised from afta (e.g. green et al. 2004; japsen et al. 2007a; turner et al. 2008). nevertheless, having a direct constraint on the palaeogeothermal gradient at the palaeothermal maximum can provide unique insights into mechanisms of heating and cooling (bray et al. 1992). 4.3.3 elevated heat flow on continental margins the approach outlined above has provided abundant evidence of significantly elevated palaeogeothermal gradients in a wide variety of settings, e.g. south-east australia (duddy 1994 1997; green et al. 2004), nw england (green 2002) and west africa (bray et al. 2002; turner et al. 2008), implying elevated basal heat flow. in contrast, many apatite fission-track studies, especially of continental margins, are characterised by the explicit assumption that palaeogeothermal gradients do not change appreciably through time, and that denudation is the only process capable of producing significant cooling in such areas (e.g. gallagher & brown 1999a, b; brown et al. 2000; gunnell 2000). this assumption is clearly erroneous, and can lead to significant overestimation of amounts of denudation. in addition to changes in basal heat flow, transport of heat due to hot fluid circulation is also a common feature in many areas (e.g. duddy et al. 1994, 1998). gallagher et al. (2005) suggested a methodology for providing rigorous constraints on palaeogeothermal gradients and amounts of removed section in similar fashion to bray et al. (1992). however, after applying such methods in namibia, raab et al. (2005) still assumed a constant palaeogeothermal gradient, despite a wide range of allowed values. as discussed in a later section, we suggest that failure to allow for elevated palaeogeothermal gradients has contributed to the inconsistency between denudation histories based on thermochronology and ideas on landscape development in the past. fig. 31 estimation of the amount of removed section from palaeothermal data. a: determination of palaeotemperatures over a depth range allows definition of the range of palaeo-geothermal gradients. extrapolation of the allowed range of palaeo-geothermal gradients then allows estimation of the amount of additional section required to explain the palaeotemperatures (equal to the amount of section that has been removed by erosion). b: higher palaeo-geothermal gradients within the allowed range will require lower amounts of removed section, and lower palaeo-gradients correspond to higher amounts of removed section. statistical procedures provide definition of the allowed range of both parameters at the 95% confidence level, as shown. note that this analysis assumes that the thermal conductivity of the eroded sequences was the same as that of the preserved sequences. vr: vitrinite reflectance. modified from green et al. (2002). d ep th , z d ril le d s ec tio n ts 0 a b maximum palaeotemperature profile present-day temperature profile to ts present-day surface temperature to palaeosurface temperature ti palaeotemperature intercept at unconformity (dt/dz) palaeogeothermal gradient palaeotemperature from afta palaeotemperature from vr re m ov ed se ct io n, z r temp., t palaeosurface re m ov ed se ct io n (k m ) palaeogeothermal gradient (°c/km) maximum likelihood 0.5 1.0 1.5 2.0 2.5 3.0 0.0 parameter range from afta, 95% confidence limits unconformity 10 20 30 40 50 60 70 ti (ti to) (dt/dz)zr = 57 4.4 other methods for constraining removed section and denudation histories 4.4.1 introduction while many studies are based on application of apatite fission-track data in isolation, significant advantages can be obtained by combining afta data with results from other methods. one reason for this is the inherent redundancy in the method (fig. 24), in that a large number of thermal histories can result in very similar afta parameters. in addition, the natural spread in the track-length distribution imposes limits on the recognition of low temperature events from afta, and can cause problems in resolving complex histories involving multiple cooling episodes. integration with independent techniques not only provides corroboration of conclusions derived from afta, but can also refine the range of thermal history solutions defined from afta alone and provide a coherent thermal history framework. here we discuss some of the most frequently used techniques that have been used together with afta and discuss the benefits that can be obtained. 4.4.2 vitrinite reflectance vitrinite reflectance (vr), based on the increase in reflectivity of the organic maceral vitrinite (a key constituent of coal) with temperature, is the standard measure of organic maturity for hydrocarbon exploration (e.g. tissot & welte 1984). the kinetics of this process are well understood (burnham & sweeney 1989), and are very similar to those of fission-track annealing in apatite (duddy et al. 1994, 1998), with vr values of 0.6% to 0.7% corresponding to total annealing of fission-tracks in typical apatites (duddy et al. 1994). these factors make vr an ideal complement to afta data applied to sedimentary sequences, as demonstrated in a wide range of studies (e.g. duddy 1997; green et al. 2004; japsen et al. 2005, 2007a, 2012b; turner et al. 2008). integration of vr data from fine-grained units with afta data from sandstones also allows determination of palaeotemperatures over a wider range of depths than possible from afta alone. the combination of both techniques can provide much tighter control on palaeogeothermal gradients and amounts of removed section than would be possible from either technique on its own. in addition, integration of vr with afta data can be of great assistance in confirming earlier events soon after deposition, which may not be confidently defined from afta alone (e.g. green et al. 2004). 4.4.3 zircon fission-track analysis and zircon (u-th)/he dating zircon is another common uranium-bearing detrital mineral amenable to fission-track analysis. from both laboratory annealing studies (tagami et al. 1996) and geological evidence (hurford 1986), fission-tracks in zirpalaeogeothermal gradients in the removed section er od ed se ct io n fo r 1 0° c p ala eo -s ur fac e te m pe ra tu re d ril le d se ct io n -1 -2 palaeogeothermal gradient in preserved section 25°c/km -3 10 °c p ala eo su rfa ce te m pe ra tu re 60 °c p ala eo te m pe ra tu re at u nc on fo rm ity su rfa ce 1501005010 palaeotemperature (°c) same lithology mix as the preserved section shale quartz sandstone salt b: 25°c/kmb unconformity a: 100°c/kma c: 16.7°c/kmc d: <12.5°c/kmd 1 0 2 4 elevation (km) 3 fig. 32. influence of the lithology of the missing section on estimation of removed thickness. differences in thermal conductivity between the removed section and the (deeper) preserved section can produce significant non-linearity in the palaeotemperature profile, which will result in unreliable estimation of the amount of removed section using the construction shown in fig. 31. only where the removed and preserved sections are identical will the thermal gradient be the same throughout the entire section. however, in practice the assumption of linearity appears to give reliable results in a wide variety of situations. vr: vitrinite reflectants. modified from green et al. (2002). 5858 con are known to be more resistant to annealing than fission-tracks in apatite, with a closure temperature around 240°c (hurford 1986) being generally accepted. investigation of zircon data in samples with different levels of vr suggests significant fission-track age reduction only occurs in zircon at maximum palaeotemperatures in excess of 250°c (p.f. green, unpublished data). therefore zircon fission-track analysis (zfta) can provide information only where temperatures of this order have been reached, and are of less relevance to landscape studies than afta and apatite (u-th)/he dating, although zfta data can be useful for providing information on sediment provenance. the general applicability of (u-th)/he dating of zircon remains uncertain at present. initial studies of this technique suggested a thermal sensitivity intermediate between afta and zfta (reiners 2005). more recent studies (guenther et al. 2013) have shown that radiation damage has a profound effect on the he retentivity, which may bring the sensitivity of low-u and high u zircons closer to that of afta while zircons with intermediate uranium contents have sensitivities closer to zfta. however, this model remains largely untested with field data. thus, while this technique has the potential to be useful in future studies of landscape development, the quantitative response of the technique remains uncertain and further work is required in order to fully understand the quantitative thermal response of (u-th)/ he ages in zircon. 4.4.4 estimating palaeoburial using sonic velocity data the progressive compaction of sediments with increasing burial has been widely used to determine former burial depths in exhumed basins (e.g. marie 1975; bulat & stoker 1987; hillis 1995; japsen 1998, 2000). comparison of compaction proxies, such as sonic velocity, in an exhumed formation with a reference curve defining the expected variation with depth in sequences of similar lithology at maximum burial depth, provides an indication of the amount of net exhumation. selection of appropriate reference curves has been problematical in some areas, leading to erroneous conclusions regarding the extent and magnitude of exhumation (cope 1986). more rigorous definition of the necessary reference curves for specific lithologies in recent years (japsen et al. 2007b) allows more reliable estimation of former burial depths, and integration of such data with constraints from afta and vr data has provided consistent reconstructions of eroded section in different settings, including the north sea basin (japsen et al. 2007a) cardigan bay, western uk (holford et al. 2005), brazil (japsen et al. 2012b) and the otway basin, south-eastern australia (tassone et al. 2013). since these compaction-based methods are controlled primarily by maximum burial depths (effective stress), in situations where they provide consistent indications of former burial depths with those derived from palaeothermal methods such as afta and vr, the results can be regarded with great confidence. in addition, while results from afta and vr can often be explained by a range of palaeogeothermal gradients and amounts of removed section, additional constraints from compaction-based methods can significantly reduce the range of viable solutions, as illustrated for example by results from the hans-1 well, offshore denmark (japsen et al. 2007a). 4.5 constraints from basic geological data and stratigraphic landscape analysis in addition to other analytical techniques, as discussed above, one form of information that is essential in ensuring that results obtained from low temperature thermochronology and these other techniques is meaningful, but is often overlooked, is basic geological data. afta data in isolation can often be explained by such a wide range of histories (e.g. fig. 24) that unless some independent constraints on the history can be defined, little or no meaningful thermal history information can be obtained. for sub-surface samples, interpretation of afta data begins by construction of a default thermal history (section 4.1.4), which is the history that can be reconstructed based on the preserved stratigraphic section and present-day thermal regime. this provides a context within which information from techniques such as afta and vr can be assessed (see e.g. green et al. 2004; japsen et al. 2007a for more details). for outcrop samples of sedimentary rock, two points of the temperature–time history provide a framework within which the post-depositional history can be evaluated: the depositional age (a time at which the apatite was demonstrably at the surface) and the present-day temperature. similar principles apply to basement immediately underlying sedimentary units. while this may seem straightforward, several ex59 amples have been cited earlier (section 4.1.5) where such basic constraints have been ignored. information provided by sla (chapter 3) can also be used to aid in interpretations of afta data, for example by recognising the presence of re-exposed peneplains indicating prior exposure at the surface (section 8.2), generations of incised valleys which may indicate staged uplift (section 5.5) and/or near-horizontal peneplains that can be used as stratigraphic markers (section 6.4.4). 4.6 summary derivation of thermal history information from fission-tracks in apatite, as well as other methods of lowtemperature thermochronology, is not straightforward. different approaches have been developed, each with their own advantages and limitations. but whichever approach is adopted, only the broad form of the history can be defined, and the degree to which the temperature–time, or t(t), solution approaches the real history depends on the complexity of the history. it is therefore essential to appreciate the limitations of the technique. interpretations of data in basement outcrop samples which assume monotonic cooling are problematical because the high degree of redundancy in the data results in a wide range of allowed histories. such histories also often conflict with geological evidence. basement samples overlain by or adjacent to sedimentary cover (however thin it may be) can provide a framework defined by geological constraints, resulting in more reliable histories. evidence from sla for periods of denudation and preservation of surfaces (by burial) can provide further constraints. results from a wide range of settings provide consistent evidence of multiple episodes of burial and exhumation (heating and cooling), and we suggest that this type of scenario provides a more suitable paradigm for interpreting data. while constraining thermal histories is difficult, defining denudation or exhumation histories (effectively palaeodepth through time) is even more challenging. in addition to uncertainty in thermal histories discussed above, factors such as variation in basal heat flow through time and other forms of heat transport (especially hot fluids, igneous intrusions) introduce further complications (although application of multiple techniques can assist to some degree). because thermal history solutions are dominated by certain key points in the history (fig. 24), the same is true of denudation/exhumation histories (fig. 33), and outside these key moments the denudation/ exhumation history is only very poorly defined. for this reason, long-term denudation rates derived from such data by averaging over long-term cooling histories have little meaning. time (ma) 050100150200250 0.5 1.0 1.5 2.0 2.5 3.0 0.0 d ep th (k m ) jurassic cretaceous cenozoictriassic modelled denudation path from aft data alternative denudation paths compatible with aft data fig. 33. monotonic vs. episodic cooling histories. given the high degree of redundancy in apatite fission-track (aft) data, as illustrated in fig. 24, it is not possible to discriminate between the continuous denudation history, shown in red, from possible alternative histories such as those shown in blue and green, involving episodic burial and exhumation. for this reason, average denudation rates integrated over the entire history have very little meaning, and the data may equally well be explained in terms of a number of episodes involving much more rapid denudation in each. 6060 5. the west greenland margin; a consistent synthesis of geological data, stratigraphic landscape analysis and low-temperature thermochronology 5.1 an integrated approach the west greenland continental margin shares characteristics with elevated, passive continental margins around the world, with asymmetrical highs reaching 1–2 km a.s.l., steeper on the oceanward side with incised valleys, sloping more gently inland, and topped with lowrelief planation surfaces (figs 34, 35). a key aspect of the region that allows a wide variety of methods to be employed in reconstructing the development of the margin (in terms of the history of subsidence and uplift) is the presence of cretaceous to eocene sedimentary and volcanic sequences exposed in the onshore nuussuaq basin (fig. 36). the evidence from these sequences themselves, and their relationship to basement rocks and landscapes developed in basement rocks, provides major insight into events during and after rifting that is not available on many other margins where cover rocks are absent. in this chapter we illustrate how an integrated approach, in which landscape analysis and palaeothermal methods are combined with geological evidence, has led to a consistent regional synthesis. 5.2 geological background: the nuussuaq basin and the offshore record 5.2.1 onshore exposures of the nuussuaq basin sediments and volcanic rocks of the nuussuaq basin are exposed in mountains capped by plateaux reaching up to 2 km a.s.l. on the peninsulas of nuussuaq and svartenhuk and the island of disko (figs 34, 36). chalmers et al. (1999) described the structural development of the basin (both onshore and offshore) and provided a summary of its stratigraphic history. figure 37 shows a summary of the stratigraphy of the basin presented by dam et al. (2009) and fig. 38 shows a summary of the subsidence and uplift history of the basin, based on the exposed stratigraphy. figure 39 shows examples of geological evidence for subsidence and uplift in the nuussuaq basin. upper cretaceous fluvial and deltaic sediments dip east (tectonic stratigraphic sequence 2, tss2; dam et al. 2009), due to rotation of originally flat-lying sediments within fault blocks (figs 39a, b, c). these deltaic sediments are unconformably overlain by latest maastrichtian and danian sediments (tss4–6, fig. 39a) and are in turn overlain by paleocene picritic and basaltic hyaloclastites and lavas (tss7, figs 39c, d, tss6 not visible; clarke & pedersen 1976; chalmers et al. 1999; dam et al. 2009; henriksen et al. 2009). the youngest rocks surviving within the onshore part of the basin are eocene lavas (storey et al. 1998; schmidt et al. 2005) below an extensive planation surface that can be traced bevelling rocks from within the nuussuaq basin into adjacent basement terrain over distances of several hundreds of kilometres (figs 40, 41; bonow et al. 2006a, b). rifting had started in the basin by the albian (the kome formation, tss1; dam et al. 2009), although reflection seismic evidence (chalmers et al. 1999) suggests that there may be older sediments than this deep in the basin. the rifting was followed by a period of thermal subsidence and the deltaic atane formation (tss2) was deposited into the accommodation space formed (figs 37, 39a, b, c). a period of tectonism and uplift occurred in the campanian (dam et al. 2009), but no equivalent event is known from the labrador sea and the origin of this episode is not understood. the eastern margin of the basin is defined by a major fault on nuussuaq. basement rocks are exposed on one side of the fault more than a kilometre above the top of cenomanian sediments, which underlie covering basalts on the other side of the fault. pulvertaft (1989) has shown that the cenomanian sediments were laid down in a low-energy fluviatile environment that shows no evidence of significant nearby topography, meaning that the fault-scarp did not exist at that time. on the other hand, farther north, mid-paleocene lavas pass across the fault without deflection by it, showing that all the kilometre-scale movement of the fault must have taken place between the cenomanian and the mid-paleocene. 61 0 0.5 1.0 1.5 2.0 elevation (km a.s.l.) upper planation surface lower planation surface rift fault, c. 65 ma paleocene–eocene basalt (onshore/offshore) cretaceous–paleocene sediment (onshore/offshore) precambrian basement (onshore/offshore) fault breakline lineament neogene afta sample(s) 1 km sukkertoppen iskappe fig. 34. topography (a) and geology (b) of central west greenland. also shown on b are topographical profiles with interpretation of two planation surfaces (upper planation surface, ups (blue lines) and lower planation surface, lps (red lines)). these surfaces are generally inclined to the east but also to the west on western nuussuaq and to the north, south of disko bugt (bonow et al. 2006a, b). they are separated in the highest areas by up to 1 km in the west but merge in the east to one surface. the surfaces cut across both precambrian basement and mid-eocene volcanic rocks and are therefore of post-mid-eocene age. mapping of the tilted and broken planation surfaces led to the identification of faults formed in connection with the uplift and tilts of the planation surfaces. three significant faults and breaklines (changes in slope gradient) relative to the planation surfaces are (1) the n–s kuugannguaq–qunnilik (k–q) fault on disko and nuussuaq, (2) an e–w fault just north of aasiaat (aa) where orthogneisses are separated from supracrustal rocks to the north; this fault separates the south-dipping planation surface on disko from the north-dipping surface south of disko bugt, and (3) the e–w sisimiut line (sl) that coincides with the precambrian ikertôq thrust zone. es depicts a hilly etch-surface, re-exposed mainly from below cretaceous cover rocks but also from paleocene volcanic rocks on southern disko and northern nuussuaq (see bonow 2005). bb, ls (on insert map): baffin bay and labrador sea. from japsen et al. 2006. 6262 fig. 35. the elevated plateau landscape of the upper planation surface in west greenland at about 900 m a.s.l. a: southern disko across paleocene basalts and b: northeast of sukkertoppen iskappe across precambrian basement. glacial reshaping has mainly affected the incised valleys, which have been widened and deepened. locations shown in fig. 36. photos: n. nielsen, university of copenhagen (a) and k. secher (b). from japsen et al. 2006. 63 nuussuaq svartenhuk halvø vaigat disko d g r disko bugt hareøen itil li f au lt ? ? palaeogene intrusive complex lower palaeogene basalt maastrichtian–paleocene sediment albian–campanian sediment precambrian basement pre-volcanic fault well location landward limit of palaeogene–neogene sediments photo location fault with lateral or alternating displacements extensional fault kq gro-3 fig. 39 c,d b a fig. 45a figs 35a, 55b fig. 41 gane-1 d a v i s s t r a i t sisimiut basin 50 km sukkertoppen iskappe 66°n 68°n 70°n 72°n 51°w54°w fig. 45c fig. 45d fig. 35b fig. 36. geological map of central west greenland. the cretaceous–palaeogene nuussuaq basin includes onshore and offshore areas from svartenhuk halvø to disko bugt. dgr: disko gneiss ridge; k–q: kuugannguaq–qunnilik fault. modified from chalmers et al. (1999), chalmers & pulvertaft (2001), bonow et al. (2007b) and dam et al. (2009). 6464 fig. 37. stratigraphy and tectonic events along a w–e line through the central part of the nuussuaq basin, west greenland, corresponding roughly to the dip direction produced by the late cretaceous faulting. note that the impact of the plume head in the danian resulted in both uplift and subsidence. tectonic stratigraphic sequences (tss) and formation names (central part of scheme) defined by dam et al. (2009). arrows indicate episodes of uplift (red) and subsidence (blue). modified from storey et al. 1998, dam et al. (2009), l.m. larsen (personal communication 2013) and g. k. pedersen & h. nøhr-hansen (personal communication 2013). kome slibestensfjeldet atane itilli kanísut mb eqalulik vaigat atanikerluk westernmost nuussuaq west nuussuaq north nuussuaq central nuussuaq south nuussuaq north-east nuussuaqtss age (ma) rift rift post-rift subsidence post-rift subsidence rift? post-rift subsidence plume-head impact ? ? ? ? ? ? volcanism volcanism uplift/ subsidence tectonism d rif t deltaic estuarine marine deep marine lacustrine fluvial hiatus subaerial volcanics maligât atane kangilia 100 110 80 90 60 70 stratigraphy of the central part of the nuussuaq basin 8 7 6 5 4 3 2 1 cenomanian turonian coniacian santonian campanian maastrichtian danian selandian thanetian ypresian albian pa le oc en e eo ce ne u pp er lo w er c re ta ce ou s pa lae og en e agatdalagatdal quikavsak 65 two phases of channel incision have been observed beneath the flood basalts in the nuussuaq basin (dam & sønderholm 1994, 1998; dam et al. 1998; dam & nøhrhansen 2001; dam 2002; dam et al. 2009). the earliest channels form part of the latest maastrichtian unconformity, and dam et al. (1998) suggested that they were eroded as submarine canyons on the crests of the fault blocks that rotated the cretaceous sediments to dip eastwards. this movement thus defines the timing of latest rifting in the nuussuaq basin close to the cretaceous– cenozoic boundary, in agreement with the known constraints on the timing of the latest rifting offshore (aram 1999; chalmers & pulvertaft 2001; christiansen et al. 2001; dalhoff et al. 2003; chalmers 2012). post-rift danian subsidence and deposition of the kangilia formation (tss4) was followed by uplift during the mid-paleocene, probably in response to the impact of the head of the iceland plume (dam et al. 1998), when new channels were eroded as subaerial valleys, which are still filled with fluviatile paleocene sediments (quikavsak formation, tss5, 6, fig. 39b). volcanism commenced shortly afterwards, during chron 27n (riisager & abrahamsen 1999), corresponding to the interval 61.65 to 61.98 ma (gradstein et al. 2004). volcanism started on a submarine slope to the west of the area where the major channels were eroded. the volcanic pile (tss7) was built up above sea level so that lavas began to be erupted subaerially. however, the lavas that entered the sea farther east formed eastward-prograding gilbert-type delta structures with cross-bedded hyaloclastite sets up to 600 or 700 m thick (fig. 39c; pedersen et al. 1993), indicating that the basin had subsided by at least this amount during the few hundred thousand base tss3 base tss4 base tss5 base tss7 surface base tss1 2 0 -2 -4 -6 -8 120 100 80 60 a 1 0 -1 -2 el ev at io n (k m a .s. l.) el ev at io n (k m a .s. l.) 70 65 age (ma) 60 b b footwall uplift thermal subsidence plume-related uplift subs. during sfs age (ma) fig. 38. summary of the subsidence and uplift history of the nuussuaq basin. a: schematic diagram showing estimates of uplift and subsidence derived from the geological record that affected the nuussuaq basin between early cretaceous rifting and shortly after mid-paleocene break-up (fig. 37). b: enlargement of part of a. the record is from dam et al. (2009) except where otherwise cited and refers to an area on the south coast of nuussuaq marked on fig. 36 (photos in fig. 39). an episode of rifting took place probably in the aptian (depositing tss1) followed by thermal subsidence that started in the albian (chalmers et al. 1999). the atane formation (tss2) was deposited during the latter event. it is at least 3 km thick but a reflection seismic line on the south coast of nuussuaq (chalmers et al. 1999, fig. 9) shows that the sediments at this location are at least 6 km thick and may be 8 km thick. an episode of uplift followed by renewed subsidence and deposition of the itilli formation (tss3) took place in the campanian. faulting, uplift and erosion of major submarine channels in the maastrichtian led to the complete erosion of tss3 at some localities such as the one represented here. the kangilia formation (tss4) was deposited in the renewed subsidence after this event (fig. 39a). two episodes in the late danian uplifted the surface to above sea level and the fluvial quikavsak formation (tss5) was deposited in the river channels (fig. 39b). submarine volcanism commenced in the western nuussuaq basin at this time. the volcanism built up a volcanic island and became subaerial at the same time as the eastern nuussuaq basin subsided by 600–700 m. subaerial lava flows from the west flowed into this basin (tss7), filling it as a hyaloclastite delta (fig. 39c). subsidence continued, however, as shown by the deposition of a succession of lava flows with hyaloclastite bases and subaerial tops (fig. 39d) and eventually entirely sub-aerial lavas. deposition of a later (eocene) subaerial volcanic succession (tss8) followed a hiatus in the volcanism. arrows indicate episodes of uplift (red) and subsidence (blue). subs. during sfs: subsidence during sea-floor spreading. tss: tectonic stratigraphic sequences in fig. 37 (dam et al. 2009). 6666 tss 4 kangilia fm tss 2 atane fm tss 5s 5 tss 5 quikavsak fm silltss 7 vaigat fmtss 7 vaigat fm tss 2 atane fm tss 2 atane fm w e 600– 700 m a c b c. 200 m hy hy hy hy hy lava lhz l l l l d tss 7 maligât fm tss 7 maligât fm tss 7 vaigat fm tss 7 vaigat fm 67 years after deposition of the fluviatile sediments of the quikavsak formation. the basin continued to subside during subsequent deposition of a 200 m thick succession of subaerial lavas alternating with five horizons of foreset-bedded hyaloclastites on nuussuaq (fig. 39d; pedersen et al. 2002), a succession recording an almost perfect balance between aggradation of the lava pile and subsidence of the basin. this succession is overlain by a 160–180 m thick zone of subaerial lavas across which the transition from chron 27n to 26r is recorded (riisager & abrahamsen 1999; pedersen et al. 2002), and above which piasecki et al. (1992) found marine dinoflagellates at a present-day height of 1176 m a.s.l., indicating subsidence was followed by substantial uplift after continental break-up in the labrador sea during chron 27 (mid-paleocene; chalmers & laursen 1995). thus the uppermost maastrichtian and paleocene sedimentary and volcanic rocks in the nuussuaq basin record uplift events that dam et al. (1998) estimated had removed up to 1.3 km of cretaceous section (fig. 39b), followed by subsidence of at least 1000 m and probably considerably more than that (figs 39c, d). at least 900 m of this subsidence took place during chron 27n (pedersen et al. 2002), which lasted 330 ka, so the minimum subsidence rate was 3 km/myr. the presence of submarine rocks today at heights of more than a kilometre above sea level shows that the events that lifted them to those heights must have taken place after the paleocene (fig. 38). the stratigraphic record onshore does not record those events, but the stratigraphic record offshore does give some clues. 5.2.2 the geological record off southern west greenland dalhoff et al. (2003) reported the presence of a major unconformity that separates mid-paleocene and younger sediments from campanian and older sediments over the whole of the sedimentary basins offshore southern west greenland (figs 42–44). they attributed this unconformity to uplift and erosion caused by impact of the head of the iceland plume, because renewed sedimentation onto the unconformity (and presumably therefore renewed subsidence) started contemporaneously with the onset of volcanism in the nuussuaq basin. the formation of the base cenozoic unconformity offshore was followed by subsidence that created sufficient accommodation space in the sisimiut basin (fig. 36) to contain a 2.5 km thickness of sediment deposited during 21 million years of late paleocene to mid-eocene time (fig. 42b; dalhoff et al. 2003). seismic sequence analysis calibrated by borehole control shows that sedimentation in the northern part of the basin was constantly sufficient to facing page: fig. 39. examples of geological evidence for subsidence and uplift in the nuussuaq basin. see fig. 37 for stratigraphy and dam et al. (2009) for stratigraphic details. photos from the south coast of nuussuaq (locations on fig. 36). a: uplift and incision during maastrichtian rifting (tss4). the fluvial and deltaic sediments of tss2 (atane fm; dam et al. 2009; also shown in b and c) were probably laid down in a basin subsiding thermally after a rift phase in the aptian–albian (chalmers et al. 1999; dam et al. 1998, 2009). the unconformity and deep channel were formed by erosion and listric faulting during the maastrichtian that appears to have removed the whole of tss3 (itilli fm), perhaps 250 m thick at this location, plus some of tss2 (atane fm). this tectonic episode affected much of southern west greenland (figs 42, 43). the channel and unconformity were buried below transgressive marine sediments of tss4 (kangilia fm), indicating renewed subsidence. there is a slight glimpse of the effects of the next phases of uplift (tss6 and tss7, volcanic rocks), but they are very foreshortened from this viewpoint. to see this mountainside from a much higher viewpoint, see dam et al. (2009, fig. 14). photo: t.c.r. pulvertaft. b: uplift and incision due to midpaleocene plume impact (tss5, tss6). renewed uplift in two phases (tss5 and tss6) during the mid-paleocene, uplifted and eroded tss4 (kangilia fm), forming new unconformities that separate the fluvial sediments of tss5 (quikavsak fm) from tss2. the quikavsak rivers flowed westwards into the sea where western nuussuaq is today. this phase of erosion is widespread over the whole of offshore southern west greenland (dalhoff et al. 2003; japsen et al. 2005) and dam et al. (1998) suggested that it is due to uplift from the rise and impact of the iceland plume head onto the base of the lithosphere. c: volcanism following early post-plume subsidence (tss7). submarine volcanism built up a volcanic island west of the present-day nuussuaq peninsula. by that time, 600–700 m of subsidence had taken place farther east, sufficient to form a marine basin into which the peridotitic lavas flowed as and eastward prograding, hyaloclastite gilbert delta in the vaigat fm of tss7 (darker volcanic rocks). dashed lines show examples of hyaloclastite foresets. the succeeding lighter volcanic rocks are the maligât fm. photo: l.m. larsen. d: volcanism during continued subsidence. subsidence continued in the late paleocene after the marine basin (shown in c) was full, shown by the 200 m thick succession of lava flows with hyaloclastite bases (hy) and subaerial tops (l), indicating that subsidence and magma production were keeping pace. eventually, either subsidence slowed or lava production increased until the volcanic succession became sub-aerial (lhz). photo and interpretation: a.k. pedersen, university of copenhagen. tss: tectonic stratigraphic sequences in fig. 37 (dam et al. 2009). 6868 fill the available accommodation space, indicated by the continuous deposition of delta-top facies. this indicates an average subsidence rate of 120 m/myr. thus the sedimentary and volcanic record preserved both onshore and offshore west greenland shows evidence for substantial mid-paleocene uplift that was probably associated with the impact of the icelandic plume, but the record also indicates that this uplift was short-lived and was followed by rapid and substantial subsidence. this uplift event could not, therefore, have been what formed the present-day mountains in west greenland. that event must have taken place later than the mid-palaeocene break-up. how and when this uplift and tilting took place has been analysed by integrating landscape analysis and afta, as described in the following sections. 5.3 results of sla in west greenland by identifying and mapping re-exposed peneplains (subcretaceous and sub-paleocene etch surfaces; i.e. hilly peneplains), high-level flat peneplains (planation surfaces), and a deep valley generation in west greenland and combining these observations with the stratigraphic record, bonow et al. (2006a, b) defined a relative denudation chronology for the development of the west greenland margin. a sub-cretaceous etch surface (es) formed by kaolinitic clay weathering of gneiss is re-exposed from below cretaceous sedimentary strata and a sub-paleocene es from below volcanic rocks on disko and nuussuaq (pulvertaft & larsen 2002; bonow 2005). the weathering products below the paleocene cover on southern disko, the saprolites, are less mature than those below the cretaceous cover, but the hills are of about the same height as those on the sub-cretaceous es. it seems thereups lps a b disko uummannaq disko bugt vaigat nuussuaq uummannaq fig. 40. the upper planation surface (ups) and the lower planation surface (lps) developed across precambrian basement (foreground) and paleocene basalt (background) on nuussuaq (view from east towards west). © danish geodata agency. modified from bonow et al. (2006a) and green et al. (2011). 69 fore likely that the old kaolinitic saprolite was stripped, and that the present relief belongs to a common late mesozoic es (fig. 45a, b). the hills are limited by fracture systems trending nw–se and by ene–wsw schistocity. these lineaments are all affected by the deep weathering. another n–s fracture system probably formed later than the es as those fractures are not affected by the deep weathering (bonow 2005). bonow (2005) used reflection seismic lines to identify the hilly surface at the sea bed south of disko and found that the surface continues over the exposed basement areas south of disko bugt (fig. 45c). the es is tilted, rising from the coast with higher elevations to the south and east. the inclined es is cut off by a younger peneplain (fig. 5), which rises slowly towards the south, where it gradually splits into two distinct peneplains, which we refer to as the upper and lower planation surface (ups and lps, respectively; fig. 34). bonow et al. (2006a, b) traced the ups across volcanic rocks as young as mid-eocene (39 ma; schmidt et al. 2005) on nuussuaq and could thus conclude that the ups is of post-eocene age. the lps is developed as a system of wide and shallow valleys cut below the ups; e.g. in the area north-east of sukkertoppen iskappe (figs 45d–f), but merges with the ups; e.g. south of disko bugt and on western nuussuaq. both surfaces are of regional extent (>500 km; figs 34, 40). the ups can be traced across the entire area. since both surfaces are continuous across rocks of different resistance to erosion, and since no single geological surface exists to which these surfaces could have been graded, the most likely explanation is that they were each graded to a general base level at the time of their formation. as the coast is known to have been near the present onshore areas throughout the cenozoic (nielsen et al. 2001; dalhoff et al. 2003), the most likely base level is sea level. a lower valley generation, glacially reshaped to a considerable extent, is deeply incised below the lps (figs 45e, f). as both the ups and lps have lost their original base of formation, they must have been uplifted and are now palaeosurfaces. the ups forms the summit surfaces, from just above sea level on western nuussuaq to over 2000 m, of crustal blocks that have been uplifted and tilted by different amounts (figs 34, 41), indicating a tectonic element in fig. 41. a: well-preserved upper planation surface (ups) across steeply dipping, palaeogene volcanic rocks on westernmost nuussuaq (view from north towards south). b: palaeosurfaces and geology along a profile (dashed line in a). note the very recent fluvial incision. the low position has preserved the planation surface from being destroyed by glaciers and cirques. location shown in fig. 36. volcanic stratigraphy in fig. 37. it: itilli fault. k–q: kuugannguaq–qunnilik fault. lps: lower planation surface. m: marine sediment (piasecki et al. 1992). © danish geodata agency. modified from bonow et al. (2006b). el ev at io n (k m a .s. l.)nuussuaq ups lps ups itk–q it 0 1 2 disko hareøen (foreground) c.1 km e w marine sediment kanísut mbmaligât fmvaigat fmcretaceous–paleocene sedimentm m n a b 7070 mid-cretaceous unconformity base cenozoic unconformity top palaeogene volcanic rocks mid-eocene unconformity oligocene–miocene unconformity early pliocene unconformity late pliocene unconformity w e w e 1 0 2 3 4 tw t (s ec ) 1 0 2 3 4 tw t (s ec ) tw t (s ec ) k pg v pg mi plio k pg v pg mi plio plio–pleisplio–pleis hs/90a-3 b 25 km sw nwikermiut-1 mi pg k plio mi pg k ggu/90-3 c d 1 0 2 3 4 5 plio–pleis plio plio–pleis a plio–pleisplio–pleisplioplio mimi pgpg pg vpg v ggu/95-14 71 the uplift (bonow et al. 2006a, b). the occurrence of marine paleocene sediments 1200 m a.s.l. (piasecki et al. 1992) close to the ups on western nuussuaq is consistent with the elevation of the ups being a measure of the amount of rock uplift since the formation of the ups at this location. the initial uplift raised the ups, which was simultaneously broken along faults, thereby forming different tectonic blocks. this phase triggered incision of new fluvial valleys to a base at sea level that ultimately led to the formation of the partly developed lps. the floors of these valleys are preserved today at a maximum elevation of 1 km a.s.l. and at a lowest level of 400 m a.s.l., revealing a second phase of differential uplift. the two uplift phases resulted in three central high areas within the study area; nuussuaq and disko <2000 m, east of nordre isortoq <1500 m and sukkertoppen iskappe <1800 m a.s.l. (fig. 34). removal of cover rocks resulted in re-exposure of the es at low elevations around disko bugt after the second uplift. bonow (2004) used geology and maps of elevation contours and slopes (fig. 46) to reconstruct a palaeodrainage system across nuussuaq and disko with a primarily southeastern direction, connected to the lower plateau remnants of the lps. the present drainage of these valleys on nuussuaq is generally to the east, but they turn westwards into the major affarsuaq valley, forming agnor (fish-hooked) valleys, indicative of an eastward movement of the water divide in connection with the glaciations. the maps of disko also indicate an originally south-easterly direction of the major valley systems. a major valley with a south-eastheading drainage has probably also occupied vaigat, the sound which today separates nuussuaq and disko. the south-east-heading palaeodrainage starts along the kuugannguaq–qunnilik fault (k–q in fig. 46a), while the palaeodrainage on the western side of this fault was towards the west. this interpretation is supported by the erosional pattern of the paleocene basalts, as areas with cretaceous and precambrian rocks exhumed from below the paleocene cover, narrow northwestwards along the suggested palaeodrainage. the headward erosion of the valleys has reached and slightly passed the kuugannguaq–qunnilik fault. the palaeodrainage helps in defining tilt direction in certain areas, e.g. in the areas between the kuugannguaq–qunnilik fault and the main cretaceous boundary fault system on nuussuaq that separates the cretaceous basin from facing page: fig. 42. reflection seismic lines from offshore southern and central west greenland showing examples of the major unconformities in the basin. see fig. 43 for stratigraphy and fig. 44 for locations of these lines. rifting took place prior to deposition of the kangeq formation (k) and pre-cretaceous sediments may also be present below this unconformity. see chalmers & pulvertaft (2001) and chalmers (2012) for details. the whole of the maastrichtian and most of the danian successions are missing between the kangeq and palaeogene (pg) successions (dalhoff et al. 2003; japsen et al. 2010). the pg succession consists of upper paleocene and eocene deltaic (on the seismic line in c) or pro-delta marine (on the seismic line in b) sediments (dalhoff et al. 2003). they are separated from overlying upper miocene (mi) sediments by a hiatus from which the whole of the oligocene and lower miocene successions are missing. the miocene and overlying pliocene (plio) sediments are mostly contourites and associated sediments. both are truncated by a late pliocene unconformity (or at the sea bed), after which thick wedges of coarse sediment prograded across the shelf edge. a: seismic line ggu/95-14. the line runs west from near the western end of nuussuaq. the top of the palaeogene lavas that reach an elevation over 2 km a.s.l. onshore dip westwards to more than 3 km below sea level. note that all the prepleistocene cenozoic sediments also dip west and are eroded at the sea bed just west of the coast, indicating that the eocene (pg) and miocene (mi) sedimentary strata must have extended farther east and must have covered at least parts of present-day nuussuaq. redrawn from chalmers (2000, fig. 4). b: seismic line hs/90a-3. note that the unconformity between the upper miocene (mi) and eocene (pg) sediments shows little or no angular truncation, indicating that uplift (and any erosion) during the oligocene and early miocene must have been very uniform over a large area (fig. 43). the platform area capped by palaeogene lavas at the eastern end of this line is the nukik platform. c: seismic line ggu/90-3 through the ikermiut-1 borehole. the sediments at the western end of this line have been folded in compressive movements along the ungava fault system (oakey & chalmers 2012), and those in the eastern part of this line have been uplifted and eroded after deposition of the miocene succession. d: detail of the area west of the ikermiut-1 borehole showing the oligo–miocene unconformity (red) and an earlier mid-eocene unconformity (green) are quite distinct at this location, being separated by upper eocene sediments. this latter unconformity truncates sediments folded during compressive movements along the ungava fault system in the eocene (oakey & chalmers 2012). the mideocene unconformity was dated at 45 ma (dalhoff et al. 2003) and upper eocene sediments are present between it and the oligo–miocene unconformity. the existence of these two separate unconformities was ignored by mcgregor et al. (2012), who came to the erroneous conclusion that the oligo–miocene unconformity was related to plate tectonic movements in the labrador sea and baffin bay. those movements had, however, ceased 10 million years earlier. plio–pleis: pliocene–pleistocene sediments. plio: pliocene sediments. mi: upper miocene sediments. pg: palaeogene sediments. pg v: palaeogene lava. 7272 fig. 43. stratigraphy offshore southern west greenland based on piasecki (2003) and sørensen (2006). note the base cenozoic, oligo– miocene and late cenozoic unconformities. dalhoff et al. (2003) attributed the base-cenozoic unconformity to uplift and erosion caused by impact of the head of the iceland plume, because renewed sedimentation onto the unconformity started contemporaneously with the onset of volcanism in the nuussuaq basin. a thick succession of reflections (‘deep sequence’, chalmers & pulvertaft 2001), below lower to mid-cretaceous sequences, offshore west greenland, and sea bed samples from the davis strait containing ordovician and jurassic sediments (dalhoff et al. 2006), indicate that palaeozoic to jurassic deposits may be present offshore west greenland as indicated in fig. 42 (see japsen et al. 2010). for c2–c4 see table 2. miocene n eo ge ne pa lae og en e la te ea rly c re ta ce ou s c en oz oi c pliocene pleistocene c4 uplift tectonic events c3 uplift manîtsoq formation stratigraphy offshore southern west greenland kangâmiut ikermiut kangeq appat kitsissut c2 uplift uplift/rifting rifting sag basin start of extension? ? ? subsidence subsidence subsidence/ volcanism plume impact thermal subsidence end compression on ungava faults quat. oligocene eocene paleocene maastrichtian campanian santonian coniacian turonian cenomanian albian aptian barremian hautevirian valanginian berriasian hiatus 140 120 130 100 110 80 90 60 40 20 30 50 70 0 10 age (ma) d rif t ? ? 73 the basement block (fig. 46a). here the ups cannot be identified due to glacial obliteration of the surface, but the direction of palaeodrainage suggests a tilt towards the south-east (fig. 46c). glacial erosion has to a large degree reshaped valleys and other areas in low positions along downfaulted blocks. not only disko bugt but also vaigat hosted major west-heading outlet glaciers that have widened and deepened the originally east-heading river valleys and also ultimately eroded through the high area around the kuugannguaq–qunnilik fault (fig. 36). the wide affarsuaq valley on nuussuaq was also largely deepened and widened westwards by glacial action. another area of pronounced glacial erosion is the outer, western parts of the low tectonic block between the two high areas between nordre isortoq and sukkertoppen iskappe. several areas here have been eroded below sea level. on the other hand, the plateaux inland are well preserved (fig. 47). in positions close to the coast and above 1000 m, an alpine relief has, in some places, developed due to erosion by valley glaciers and cirques (fig. 5). in the tectonic block east of the kuugannguaq–qunnilik fault on nuussuaq, the ups is totally destroyed in this way. in contrast, the ups is well preserved in low positions on westernmost nuussuaq (fig. 41). 5.4 use of palaeothermal methods to define an absolute chronology of uplift and exhumation events the results of sla presented in section 5.3 provide clear definition of multiple uplift events of west greenland. a period of prolonged uplift and erosion took place after paleocene break-up west of greenland and led to development of a regionally extensive peneplain that cuts across palaeogene basalt and older rocks. this peneplain (ups), now recognised as plateaux up to 2 km a.s.l., is broken and tilted in different directions and now extends from sea level to the highest summits. a less tilted, lower planation surface (lps) occurs between 400 and about 1000 m a.s.l. (fig. 40). the deep valleys incised below the lps are more or less glacially reshaped. thus, three episodes of uplift and erosion occurred after break-up: the first episode led to formation of the ups and two subsequent episodes lifted the ups to its present elevation. low-temperature thermochronology studies, undertaken to place these episodes into an absolute timescale (fig. 48), were described by japsen et al. (2005, 2006, 2009) fig. 44. outline of the extent of the oligocene – lower miocene hiatus offshore southern and central west greenland based on seismic and well data (sørensen 2006; see the offshore stratigraphic column in fig. 43). oligocene strata have been encountered east of labrador on the canadian side of the labrador sea. no information is available for the area between the two areas where oligocene is marked as absent and present, respectively. the oligocene hiatus thus represents the time interval during which the oligo–miocene peneplain, ups, was graded to sea level onshore after the uplift event that began at the eocene–oligocene transition. bathymetry and bedrock topography in metres. icecaps indicated as white. seismic lines and cross-sections shown in figs 42 and 53 are indicated. modified from japsen et al. (2010). 60˚w baffin island labrador 70˚n 66˚n 62˚n 58˚n 70˚w 50˚w -4000 -3000 -2000 -1000 -200 0 500 1000 2000 3000 3500 elevation (m a.s.l.) greenland oligocene absentoligocene present o nshore uplift at 35 m a fig. 42a fig. 53 fig. 42c fig. 42b qulleq-1 ikermiut-1 gro-3gro-3 200 km 7474 with detailed results presented as online supplementary data files. we present a summary of results in the following. afta data were collected in a large number of outcrop samples from the nuussuaq basin and adjacent basement regions (65–72°n), resulting in definition of a series of phanerozoic cooling episodes (table 2). one precenozoic cooling episode is of importance for this study, namely the late jurassic event which began between 160 and 150 ma (cooling episode c1). this event is indicated in afta data from most samples from the basement areas, and preceded the onset of known rifting offshore and in the nuussuaq basin (chalmers & pulvertaft 2001). the es is the probable end-result of land-form evolution following this cooling episode, which thus must have involved exhumation. a c 800 m a.s.l. m a.s.l. 700 600 500 400 300 200 100 0 b d f f f e 1600 1500 1400 1300 1200 1100 1000 900 800 700 600 sukkertoppen iskappe ka ng erl uss ua q gneiss gneiss basalt ups lps slopes lakes ice hill complex basalt 500 400 300 200 100 1 km 0 1 km fortuna bay n fig. 45. four distinct landscape types in west greenland. a, b: the re-exposed sub-paleocene etch surface on southern disko. the stippled line shows the approximate contact between paleocene basalt and precambrian gneiss. the summit plateau is the upper planation surface (ups). c: the hilly etch surface on the mainland, just south of disko bugt. the hills are up to 100 m high. d: the ups and the lower planation surface (lps), east of sukkertoppen iskappe (location shown in e). e: 3d model of the landscape east of sukkertoppen iskappe with the ups, lps and incised valleys. f: interpretation of surfaces east of sukkertoppen iskappe (location shown in e) with ups, lps and incised valleys, marked with brown slopes with more than 12° inclination. photo locations are also marked in fig. 36. from bonow et al. 2007b. 75 the most detailed constraints on the cenozoic history were obtained from analyses in and around the nuussuaq basin, where afta data were supplemented by vr data in cretaceous and palaeogene sedimentary rocks. cretaceous and palaeogene sedimentary and volcanic rocks rest directly on basement in the nuussuaq basin, and this immediately shows that the basement was at the surface at the time of deposition. these conditions provide useful geological constraints on the mesozoic– cenozoic cooling/exhumation and heating/burial history of the basement rocks. the tightest constraints on the timing and magnitude of three episodes of cenozoic cooling were obtained from afta and vr data in palaeogene and upper cretaceous sediments in two boreholes (gro-3 and gane-1) down to 3 km depth in western nuussuaq (figs 36, 49). combined with results from regional outcrop samples, these data show that the sedimentary units began to cool from their maximum post-depositional palaeotemperatures at some time in the interval 36 to 30 ma (c2, eocene–oligocene transition), with two subsequent cooling episodes beginning in the intervals 11 to 10 ma (c3, late miocene) and 7 to 2 ma (c4, latest miocene–pliocene). the eocene–oligocene cooling episode is also recognised in afta data from outcrop samples of basement and sedimentary rocks across the region, with the highest palaeotemperatures prior to the onset of cooling observed at locations at the western end of nuussuaq (fig. 50), defining a distinct ‘hot spot’ close to the location of the gro-3 borehole. this ‘hot spot’ reflects an elevated palaeogeothermal gradient at this location compared to surrounding regions. note that the two most recent episodes are only recognised in deeper afta samples which cooled from palaeotemperatures around 80°c or more in these events (fig. 51a). in surface and near-surface samples, peak temperatures in these episodes are of insufficient magnitude to be resolved in the afta data. the eocene–oligocene cooling episode started 30 million years after the final rifting which occurred around the k/t boundary, more than 25 million years later than peak volcanism during the paleocene and c. 10 million years later than the youngest recorded volcan7720 7760 7800 20 km contour interval 300 m 7840 580 620 660 700 740 k–q k–q it it cbfs cbfs ki a i av k s i av vaigat disko bugt ice palaeogene basalt fault cretaceous– paleocene sediment post-paleocene sediment precambrian basement shear zone ttilted blockfault summit (elevation in m a.s.l.) fault inferred palaeodrainage slope 0–6.5° 6.5–12° 12–25° >25° t t t t t t t 1904 1940 2144 b c 2010 fig. 46. reconstruction of the palaeodrainage for nuussuaq and disko. a: geology and relief. b: slopes. c: reconstructed palaeodrainage. av: affarsuaq valley. cbfs: cretaceous boundary fault system. it: itilli fault. ki: kingittoq fault. k–q: kuugannguaq– qunnilik fault. s: saqqaq valley. utm coordinates (km) zone 22n. from bonow 2004. 7676 ism, yet the afta data show that the upper cretaceous sedimentary units were hotter at the end of the eocene than they were during the volcanic and rifting phases. maximum palaeotemperatures derived from afta and vr data in the two boreholes are highly consistent (fig. 51a), noting that the vr data only record the maximum temperature event. the availability of the vr data provides more reliable definition of the palaeotemperature profile characterising this event than would be possible from afta alone which only provides minimum limits on the maximum palaeotemperature in deeper samples. the combined dataset for the gro-3 and gane-1 boreholes define an eocene–oligocene palaeogeothermal gradient in the range 39 to 44°c/km, and require that an additional c. 1925 m (between 1750 and 2100 m) of section must have been present above the borehole location (close to sea level) at the palaeothermal maximum (fig. 51b; japsen et al. 2005, note added in proof). some of this section certainly consisted of basalts that are still present today in the mountains around the borehole (fig 51c). however, it is unlikely that all of it was basalt and part of the former section above today’s highest plateaux may well have consisted of eocene sediments that transgressed from the west. such a section is seen on offshore seismic data, but today it is truncated by a shallow unconformity or the sea bed some distance offshore (fig. 42a; chalmers 2000). fig. 47. the well-preserved inland plateau (the upper planation surface) across precambrian basement and cut by deep valleys and fjords; view towards south-west from a position just south of sukkertoppen iskappe (fig. 34). plateaux inland are well preserved, although glacial erosion has stripped all shallow saprolites from the surfaces and lakes have formed, but no deep erosion has occurred. the photo illustrates how surfaces of resistant bedrock remain flat even after glacial erosion. elevation of the plateau about 1 km a.s.l. © danish geodata agency. 77 5.5 discussion of afta data from west greenland redfield (2010) questioned the interpretation of afta data and the supporting information from west greenland presented by japsen et al. (2005, 2006, 2009), claiming that many factors had been ignored which could have significant impact on the data. however, redfield (2010) did not mention a large number of publications in which these factors, and the way in which they are accounted for in the interpretative methods employed by japsen et al. (2005, 2006, 2009), were discussed at length (many of these publications are discussed and referenced in sections 4.1 and 4.2). for example, redfield (2010) claimed that the two neogene cooling events defined from the afta data are ‘model artefacts’. however, as discussed above these events are also clearly documented in the landscape and in the geology of the region. the timing constraints on these episodes are derived from afta data in samples from depths >1 km which show that the samples were heated to >120°c (supported by vr data). apatites in these samples have thus only retained tracks since neogene cooling. redfield (2010) accepted that regional planation and subsequent uplift/ exhumation clearly postdates eocene basalts which are truncated by the ups, and further acknowledged that neogene uplift and erosion “may well have occurred”. green et al. (2011) emphasised that their interpretation is supported by a wealth of independent information, and therefore considered their interpretation to be highly reliable. 5.6 integration of palaeothermal data with sla a major problem in previous landscape studies was the difficulty in dating the formation of epigene surfaces and fig. 48. summary of stratigraphy offshore and onshore west greenland, cooling events identified onshore from afta, peneplain formation and valley incision, and key regional tectonic events as seen in the structural systems onshore. onshore and offshore unconformities correspond to regional tectonic events that are identified by stratigraphic landscape analysis and constrained in time by the cooling events defined by afta. weathering and erosion of the etch surface (es) in basement rocks took place following late jurassic exhumation (cooling event c1, table 2) and prior to early cretaceous burial, and renewed development took place prior to paleocene volcanism. following post-breakup subsidence and burial, a first phase of uplift and exhumation that began at the eocene–oligocene transition (c2), led to the formation of the upper planation surface (ups) during the oligo–miocene (see fig. 54). this surface was offset by reactivated faults, resulting in megablocks that were tilted and uplifted to present-day altitudes of up to 2 km in two phases that began in the late miocene and in the latest miocene–pliocene, c3 and c4, respectively. the c3 uplift led to incision below the uplifted ups and thus to formation of the lower planation surface (lps). the peneplains and incised valleys thus reveal information on erosion from periods in the past from where there exist no other geological data. dotted line: maximum age range of sediments. onshore stratigraphy from fig. 37 and schmidt et al. (2005), pedersen et al. (2006). offshore stratigraphy from fig. 43. ec: early cretaceous. lc: late cretaceous. pl: plio–pleistocene. modified from bonow et al. 2007b and japsen et al. (2009). geology onshore tectonic events geology offshore peneplain and valley formation onset of cooling (afta) faults/fractures system onshore eses ups lps va wrenchingextensionextension opening r: rifting phase vo: volcanism d: drifting phase uplift event es: etch surface period of deposition onset of cooling ups: upper planation surface lps: lower planation surface va: valley formation 150 100 50 40 30 20 10 plmioceneoligoceneeocenepaleocenelcecjurassic 6070 age (ma) 0 c1 c2 c3 c4 r r dvo vo vo 7878 their uplift. this difficulty has been overcome in west greenland by combining the relative chronology derived from the landscape with application of afta and vr data to define the absolute timing and magnitude of the corresponding erosional episodes as discussed in the preceding section as illustrated in fig. 52. here we explore the relation between the three post-breakup episodes of cooling and exhumation revealed by afta data and the steps in the landscape defined by the ups and the lps. the cooling episode that began at the eocene– oligocene transition is recognised in afta results from almost every sample analysed from within the nuussuaq basin and in the surrounding basement areas (japsen et al. 2006). the samples recording this event are distributed across an area similar in extent to the ups that, as described earlier, has been dated independently to be post-mid-eocene (fig. 34). japsen et al. (2006) therefore considered that both the formation of the ups and the cooling at the eocene–oligocene transition are expressions of the same episode of exhumation. the gro-3 borehole was sited on westernmost nuussuaq in a valley bottom near sea level in a mountainous area. this allows the possibility of investigating the relation between the cooling events defined from the gro-3 data and the elevation of the ups and lps that we have mapped in this area (bonow et al. 2006b). as shown in fig. 51c, the ups forms the local summit level at a height of c. 1250 m a.s.l. (1100–1400 m a.s.l.) in the vicinity of the gro-3 borehole. at the onset of cooling from the eocene–oligocene c2 palaeothermal maximum, a cover of c. 1925 m (1750–2100 m) was present above the drill site (fig. 51b, point 1). only part of the removed cover can be explained by the basalt present in the adjacent mountains (fig. 51c) and an additional c. 675 m (350–1000 m) of section must have been present above the present-day summit level (ups) at the palaeothermal maximum. this additional section must have been deposited after the youngest preserved basalt was laid down and prior to the onset of c2 exhumation in the fig. 49. palaeothermal episodes recognised from afta data in the gro-3 and gane-1 boreholes. synthesis of all data suggests three discrete episodes of cooling (vertical bands). analysis of a larger dataset from the nuussuaq basin allows the onset of cooling episodes to be refined further to 36–30 ma (c2, eocene– oligocene transition), 11–10 ma (c3, late miocene) and 7–2 ma. (c4, latest miocene – pliocene); see table 2. modified from japsen et al. (2005). time (ma) regional constraints 36 –3 0 m a 11 –1 0 m a 7– 2 m a gc883-3 gane-1 0102030405060708090100110120130140150 gc883-8 gc883-9 gc883-10 gc883-11 gc883-12 gc883-13 gc883-14 gro-3 stratigraphic age range palaeothermal constraints from afta sample regional constraints on the onset of cooling onset of cooling (ma) stratigraphic interval cooling episode 560–500 latest neoproterozoic – cambrian 370–355 late devonian – early carboniferous 230–220 late triassic c0 160–150 late jurassic c1 36–30 eocene–oligocene transition c2 11–10 late miocene c3 7–2 latest miocene – pliocene c4 table 2. intervals for onset of cooling episodes in west greenland intervals defining the beginning of episodes of regional cooling derived from afta data in 69 samples from outcrops and boreholes in west greenland. cooling episodes after japsen et al. (2006, 2009) and bonow et al. (2007b). 79 interval 36 to 30 ma, and subsequently removed during the formation of the ups. constraints on palaeogeothermal gradients and removed section for the two more recent episodes are less well defined (fig. 51b), largely because these are only defined from afta and palaeotemperature constraints are relatively broad (fig. 51a). but as indicated by point 3 in fig. 51b, for a palaeogeothermal gradient of c. 30°c/km, the latest miocene – pliocene (c4) palaeotemperatures can be explained by around 850 m of additional section, which corresponds with the height of the lps at this location. in other words, the ‘pliocene’ cooling episode, which began between 7 and 2 ma, can be explained simply by incision and excavation of the present relief below the lps within the valley where the borehole is located. this event can therefore be interpreted as representing the onset of uplift that took the lps to its current altitude. the question then remains as to the evolution between points 1 (c2 maximum burial at c. 35 ma) and 3 (onset of latest miocene – pliocene c4 uplift and incision) in fig. 51b. one option involves a scenario in which there was no cover above the ups at c. 10 ma when the late miocene c3 cooling phase began (point 2). at point 2, the cover required above the ground surface to explain the late miocene palaeotemperatures corresponds to the present-day elevation of the ups, for a palaeogeothermal gradient of c. 40°c/km (i.e. the amount of rock still preserved above the level of the drill site in the sides of the valley). in this scenario, the 10 ma cooling episode can be explained in terms of excavation of the valley in which the well is sited, and the onset of cooling at 10 ma would therefore indicate the beginning of uplift leading to incision below the ups. with the wide range of possible combinations of removed section and palaeogeothermal gradient allowed by the late miocene palaeotemperatures, a range of alternative scenarios is possible, as shown by the dashed trajectories in fig. 51b. each of these alternative trajectories involves some degree of cover on the ups at the onset of late miocene cooling, with the precise amount depending on the value of late miocene palaeogeothermal gradient. in the previous section, we argued that the ups obtained its final shape after protracted denudation ? ? ? ? gro-3 itilli fault 72°n 71°n 70°n 69°n 55°w 50 km 53°w 51°w palaeogene intrusive complex lower palaeogene basalt maastrichtian–paleocene sediment albian–campanian sediment precambrian basement pre-volcanic fault fault with lateral or alternating displacements extensional fault <60 end-eocene palaeotemperatures (°c) 60–70 70–80 80–90 90–100 100–110 >110 fig. 50. eocene–oligocene (c2, 36–30 ma, table 2) palaeotemperatures from afta in samples from nuussuaq and surrounding regions. the map shows a pronounced ‘hot spot’ on western nuussuaq close to itilli fault (and the location of the gro-3 borehole). this reflects an elevated palaeogeothermal gradient at this location compared to surrounding regions. boreholes are represented by the shallowest sample in each case. while samples were taken from different elevations, these results have not been adjusted to a common datum. because the range of elevations is less than c. 500 m this will not significantly distort the regional pattern. values in individual samples and further discussion in supplementary data files to japsen et al. (2005, 2006). modified from green et al. (2011). 8080 surface at maximum burial (c. 35 ma) sea level removed palaeogene cover removed paleocene basalt paleocene basalt cretaceous– paleocene sediment gro-3 lps lps ~6 75 m ups ups ~1 92 5 m ~1 92 5 m ~1 25 0 m ~850 m temperature (°c) late eocene profile c2 0 20 40 60 0 20 40 palaeothermal gradient (°c/km) assumed present-day thermal gradient: 30°c/km 605030 80 100 120 140 160 180 200 220 240 260 0.0 0.5 1.0 1.5 2.0 2.5 3.0 0.0 0.5 1.0 1.5 2.0 2.5 3.0 d ep th b el ow se ale ve l ( km ) re m ov ed se ct io n (k m ) gane-1 gro-3 assumed present-day temperature profile latest m iocene– pliocene profile c4 maximum palaeotemperature from vr maximum palaeotemperature from afta late m iocene profile c3 late miocene latest miocene– pliocene 1250 m 1925 m eocene–oligocene 850 m a b c 1 3 2 fig. 51. reconstruction of the burial, uplift and exhumation history around the gro-3 borehole, western nuussuaq, based on afta and stratigraphic landscape analysis (location shown in fig. 36). a: palaeotemperatures in three episodes identified from afta and vitrinite reflectance (vr) in the gro-3 and gane-1 boreholes; eocene–oligocene (c2, onset between 36 and 30 ma), late miocene (c3, onset between 11 and 10 ma) and latest miocene – pliocene (c4, beginning between 7 and 2 ma; table 2). b: ranges of palaeogeothermal gradient and removed section required (banana-shaped areas) to explain the palaeotemperatures in each episode from a. the well-defined palaeothermal maximum (point 1) corresponds to a cover above the valley floor of c. 1925 (1750–2100) m for a palaeogeothermal gradient between 39 and 44°c/km. point 3 in b corresponds to a cover above the valley floor of c. 850 m thickness for a palaeogeothermal gradient of c. 30°c/km, corresponding with the height of the lps at this location, and thus the latest miocene – pliocene cooling episode can be explained simply by incision of the present relief below the lower planation surface (lps). several trajectories are indicated between points 1 and 3. point 2 represents a scenario in which there was no cover above the upper planation surface (ups) when the late miocene cooling phase began. at point 2, the cover required above the ground surface to explain the late miocene palaeotemperatures corresponds to the present-day elevation of the ups (c. 1250 m or 1100–1400 m a.s.l.), for a palaeogeothermal gradient of c. 40°c/km. in this scenario, the 10 ma (c3) cooling episode can be explained in terms of excavation of the valley in which the well is sited. the onset of cooling at 10 ma would therefore indicate the beginning of uplift leading to incision below the ups. the dashed lines in b indicate the wide range of possible combinations of removed section and palaeogeothermal gradient allowed by the late miocene palaeotemperatures. each of these alternative trajectories involves some degree of cover on the ups after its final formation, at the onset of late miocene cooling, with the precise amount depending on the value of late miocene palaeogeothermal gradient. c: reconstruction of the total amount of rock present above the site of the gro-3 borehole at the eocene–oligocene palaeothermal maximum relative to the elevation of the ups and lps in the adjacent mountains; c. 1250 (1100–1400) and c. 850 (800–900) m a.s.l., respectively. the amount of section present at the eocene–oligocene palaeothermal maximum extends above the local summit level (the ups) by c. 675 (350–1000) m; consequently, the formation of the ups involved removal of that section. modified from japsen et al. (2005, 2009). 81 fig. 52. conceptual model highlighting the relationships between stratigraphic landform analysis, afta and onshore/offshore geology in central west greenland. the red and the blue dots in sketch a mark two afta samples with different exhumation histories (see d and e). a: sketch of landforms in central west greenland showing four different erosion surfaces from es (oldest), ups, lps to s4 (youngest). the age of a re-exposed etch surface (es) is well constrained due to protective cover rocks. the upper and lower planation surfaces (ups and lps) cut across both strata of different age and the es. s4 is a newly formed surface. the surfaces and the relationships to geology can be directly observed in the landscape. b: a relative event chronology from the classical geomorphological interpretation of the landforms. cover rocks (k: cretaceous, p: palaeogene) constrain the timing of the final formation of surfaces (es), but may also give the maximum age of surfaces (ups). the dashed lines indicate the approximate time needed for completion of a surface. the dashed lines indicate the uncertainties for their formation in time. time line from geology and afta data. c: the sedimentary sequences along the passive margins are tilted and truncated below the plio–pleistocene strata due to late uplift of the landmass. the oldest sequence is thus exposed closest to the coast. a significant unconformity spans late eocene to late miocene. glaciers expanding to the shelves have also contributed to the erosion of offshore sequences. d: uplift and burial history of the red sample in a (exposed at es near paleocene basalt). afta data record late jurassic onset of cooling (exhumation) from palaeotemperatures of about 75°c corresponding to burial below a cover of about 2 km depending on the palaeogeothermal gradient. near sukkertoppen iskappe, palaeozoic sediments were present near the present surface in the late jurassic (japsen et al. 2009), implying that the cover at that time included palaeozoic–jurassic rocks as in the case shown here. because the es formed in exposed basement, this overburden must have been removed prior to deposition of paleocene basalts (this case) or lower cretaceous sediments (elsewhere in the nuussuaq basin) that has preserved the es at certain locations. the es was again re-exposed at the surface due to uplift and erosion, recorded as onset of cooling by afta and as truncation of the sedimentary sequences offshore. e: burial and uplift history of the blue sample in a, exhumed at the valley bottom (s4). formation of the ups included erosion of both overburden and basement, as the es was obliterated by the planation event. reburial of the ups is suggested by afta data (see fig. 51). uplift results in the development of a new planation surface, the lps by valley incision below the ups, tilted sedimentary sequences offshore and afta recording onset of cooling. further uplift raised the ups and lps to their present elevation. the fluvial v-shaped valleys were reshaped by glaciers. the development of the new s4 surface was initiated. modified from bonow et al. (2007a). eocene miocene plio-pleistocene sea es es es ups ups ups ups lps lps s4 es s4 s4 s4 k pk p lps sea sea lps s4 a b c d e time (ma) time (ma) 60 time (ma) 1535 10 5 0 160 1 3 2 1 0 depth (km) 0 1 2 3 4 35 0 100 60 10 5 0 es es ups ups afta sample es ups lpsprecambrian basement palaeozoic–jurassic cretaceous paleocene volcanics ups lpseses 8282 during oligo–miocene time, but prior to late miocene (c. 10 ma) uplift that initiated the dissection. the scenarios involving some degree of cover on the ups indicated by the dashed lines in fig. 51b thus require that sediment accumulated on the ups after its final formation and before late miocene uplift. our preferred solution involves some degree of reburial of the ups. we base this interpretation on the presence of a thick miocene sequence, tilted and truncated just offshore to the west of nuussuaq, which provides strong evidence that miocene strata once extended across at least western nuussuaq (see section 8.7). combining results from nuussuaq and the basement areas to the south demonstrates that the ups and lps were formed by denudation across the entire area of central west greenland, even in the basement area where no remnants of cover rocks are preserved. interpretation of afta data from outcrop samples up to 1.8 km a.s.l. near sukkertoppen iskappe (fig. 34) shows that the ups is the end-product of post-eocene denudation even in basement areas, and that phanerozoic sediments – most likely of cretaceous–palaeogene age – must have been present above this level prior to the onset of denudation (japsen et al. 2009). in conclusion, the results reviewed above suggest that the three post-breakup events of uplift and exhumation identified in the afta data reflect the events that led to the distinct levels in the landscape. (1) uplift and exhumation that began at the eocene–oligocene transition led to formation of the ups. (2) uplift and exhumation that began in the late miocene led to formation of the lps by incision below the uplifted ups (after removal of any miocene cover). (3) latest miocene – pliocene uplift and exhumation led to formation of the present-day fjords and valleys below the uplifted lps. this interpretation implies that the ups was graded to base level through oligocene and miocene times until the late d ep th b el ow se alev el (k m ) 1.00.50.30.2 0.70.4 qulleq-1 default history l. eocene – m. miocene unconformity recent pliocene m. – u. miocene l. eocene u. paleocene l. campanian u. santonian 1 1.5 2 2.5 3 maturity (%ro) °c/km: 20 25 30 b st9902-01001 wnw ese 2.0 2.5 1.5 3.0 3.5 4.0 qulleq-1 tw ow ay tr av el ti m e (s ec ) 2 km uc2 qm qm: quaternary – middle miocene leoc: lower eocene upal: upper paleocene uc: upper cretaceous lc: lower cretaceous lc basement upal uc1 leoc a fig. 53. the qulleq-1 well, offshore west greenland. a: seismic line through the qulleq-1 well. b: plot of vitrinite reflectance (vr, red dots) data vs. depth for the qulleq-1 well together with the drilled succession. in a, note the pronounced angular unconformity below the neogene succession that truncates lower eocene, paleocene and upper cretaceous strata. in b, black curves indicate predicted vr trends based on the default history (preserved stratigraphy and geothermal gradients ranging from 20 to 30°c/km). the discrepancy between observed and predicted vr values indicates that the section has been hotter in the past, most likely due to deeper burial below a section that was subsequently removed. only the vr value for the sample above the base-neogene unconformity matches the default history. this suggests that the early eocene – middle miocene hiatus represents the removal of sediments of that age. location shown in fig. 44. modified after christiansen et al. (2001) and japsen et al. (2010). 83 miocene, when a phase of uplift initiated the destruction of the ups by river incision. 5.7 correlation with the offshore record four sedimentary successions of cenozoic age have been recognised offshore southern west greenland (dalhoff et al. 2003; piasecki 2003; sørensen 2006): mid-paleocene to late eocene, middle to late miocene, early pliocene and late pliocene to pleistocene (figs 42, 43). these are separated by hiatuses of oligocene to early miocene, late miocene, and early pliocene age. deposition of the first part of the earliest sequence corresponds with the period of subsidence recognised in the nuussuaq basin from the exposed stratigraphy and palaeothermal data. the oligocene to middle miocene hiatus formed during the same period as the ups (beginning around the end of the eocene, c. 35 ma). there is little evidence of truncation of sedimentary reflectors at this unconformity (fig. 42), so the uplift that formed the ups seems to have been fairly uniform over a large area of the continental shelf as well as the present-day onshore (japsen et al. 2006). correlating the two most recent hiatuses with the two phases of late neogene uplift suggests that the late miocene hiatus formed during the uplift that began at c. 10 ma and resulting in formation of the lps, while the early pliocene hiatus provides a more precise timing for the onset of the final uplift phase at c. 4 ma (figs 43, 48). the total magnitude of neogene vertical motions along the coast of west greenland can thus be defined by the depth of the base miocene offshore and the elevation of the ups onshore. because of its regional importance, we will discuss the nature of the oligocene hiatus at some length in the next section. 5.7.1 the nature of the oligocene hiatus offshore southern west greenland oligocene sediments are absent on the shelf off central and southern west greenland (sørensen 2006) and, north of about 66°n, upper eocene sediments are separated from lower eocene sediments by another pronounced hiatus (figs 43, 44; dalhoff et al. 2003). south of about 66°n, the two hiatuses merge into one and upper miocene sediments lie directly on lower eocene sediments (fig. 42; cf. nøhr-hansen 2003; piasecki 2003; sønderholm et al. 2003; sørensen 2006). the earlier hiatus coincides with the substantial decrease in the speed of sea-floor spreading in the labrador sea after chron 21 and the latter hiatus with the final cessation of seafloor spreading between greenland and north america (chalmers & pulvertaft 2001). it is not possible to use only reflection seismic data to evaluate whether a hiatus represents a period of nondeposition or if sediments have been deposited and then removed. the oligocene hiatus is seen as a very lowangle unconformity on seismic data from much of the greenland shelf, and it can thus be difficult to judge if it is an erosional unconformity (figs 42a, b). such low-angle unconformities are conventionally regarded as representing intervals of stability, but their character seems to be different from the thirdand higher order unconformities studied by sequence stratigraphy. dalhoff et al.’s (2003) sequence stratigraphic analysis of the mid-paleocene to mid-eocene succession of southern west greenland identified 11 third-order sequences between the base cenozoic unconformity and the mideocene unconformity, each separated by a third-order unconformity that probably formed in response to variations in sea-level. sedimentation in at least the northern sisimiut basin was, however, continuous enough to keep the accommodation space filled. the base cenozoic, oligocene–miocene and probably the late miocene and plio–pleistocene unconformities appear to have been the result of a process that interrupted the subsidence of the basin and probably involved uplift and erosion of previously deposited sediment. evidence for this is available from the qulleq-1 exploration well. the oligocene to middle miocene unconformity in the area penetrated by the qulleq-1 exploration well is clearly angular and truncates lower eocene, paleocene and upper cretaceous strata (fig. 53a; christiansen et al. 2001). vr data from the qulleq-1 well suggest that the pre-neogene succession in the well has been more deeply buried in the past. vr data from the well (fig. 53b) are seen to plot above curves of predicted vr trends based on the ‘default history’ derived from the preserved stratigraphy and based on a likely range of geothermal gradients between 20 and 30°c/km (in this history, hiatuses are assumed to represent periods of non-deposition, not erosion). the discrepancy between observed and predicted vr values indicates that the drilled section below the unconformity has been hotter in the past, most likely due to deeper burial below a section that has been removed. 8484 a a' a''nuussuaq disko hareøen t w t ( se c) 0 1 1 2 2 a a' a'' a a' a'' a' a'' h ei gh t (k m ) sea bed base quaternary base pliocene base miocene mid-eocene unconformity upper planation surface lower planation surface etch surface uplift (miocene, plio–pleistocene) subsidence (miocene, plio–pleistocene) marine paleocene sediment fault m 0 0 0 0 0 0 700 km600 0 k–q rfga p m gro-3 160 ma 70 ma 35 ma 15 ma 11 ma present 120 ma palaeozoic–jurassic sediment precambrian basement cretaceous– paleocene sediment plio–pleistocene sediment gro-3 ka m v 20 km 5 ma palaeogene sediment paleocene–eocene basalt miocene sediment 85 this possibility is strengthened by the observation that the vr value for the only neogene sample matches the default history, indicating that this sample is now at its maximum burial depth. further studies based on palaeothermal (afta and vr) data and palaeoburial (sonic) data and constrained by the preserved stratigraphy, are needed to evaluate the timing, magnitude and extent of the exhumation related to the oligocene hiatus offshore west greenland. such studies would provide firm constraints on both the thermal and the burial history of the preserved sedimentary section, allowing definition of those areas where maximum burial was reached during mid-cenozoic times. studies of the burial and exhumation history west of greenland are also important for understanding the tectonic evolution of the arctic during the cenozoic within a wider context: a hiatus similar to that encountered on the west greenland shelf is present in the sverdrup basin in the eastern canadian arctic (harrison et al. 1999), and a similar stratigraphic break was unexpectedly penetrated on the lomonosov ridge (central arctic ocean) by the integrated ocean drilling program (iodp) expedition 302 (backman et al. 2008). the drilled core documented a 26 million years hiatus, separating middle eocene (c. 44 ma) from lower miocene sediments (c. 18 ma). sangiorgi et al. (2008) were unable to determine whether that hiatus was generated by sediment erosion or by non-deposition, but emphasised that it conflicts with classical post-rift thermal subsidence models for passive margins. the temporal correlation between this oligocene hiatus and the time interval, during which the oligo–miocene peneplain, ups, was graded to sea level onshore west greenland, emphasises the regional controls on such processes, as also highlighted by green & duddy (2010). facing page: fig. 54. reconstruction of tectonic events in central west greenland from the palaeozoic to the present-day, with focus on the late cenozoic development (cf. fig. 48). late palaeozoic to mesozoic cover was removed after late jurassic uplift and prior to development of a sub-aerial weathered etch surface (bonow 2005; japsen et al. 2006, 2009). this surface was covered by cretaceous sediments and paleocene basalts. after post-rift subsidence and renewed uplift, a regional upper planation surface (ups) developed across both basement and basalts. then followed a period of subsidence, with probable development of a miocene cover across the ups; notice the tilted and truncated miocene sequence just west of nuussuaq (fig. 42a) and the freshness of the ups on the western tip of nuussuaq (fig. 41). the subsequent uplift of the ups resulted in removal of the miocene sediments and development of the lower planation surface (lps) until about 5 ma when a last phase of uplift led to incision below the lps. twt: two-way travel time. from japsen et al. 2006. b 0 1 tw ow ay t im e (s ec ) 2 sw 20 km ne miocene quaternary paleocene–eocene a sediments deposited after phases of uplift and erosion post-rift, thermal subsidence sediment and volcanic rocks volcanics post-rift unconformity fig. 55. offshore-onshore correlation along the elevated, passive continental margin (epcm) of west greenland. a: highlighted oligo–miocene hiatus (blue dashed line) on a seismic section off nuussuaq (see fig. 42a; location shown in fig. 44). b: highlighted oligo–miocene peneplain (the upper planation surface, ups, blue dashed line) across paleocene basalts on southern disko (location shown in fig. 34). the present-day high topography of the west greenland epcm formed since the late miocene, by uplift and dissection of the ups, the remnants of which can now be recognised as plateaux at elevations up to almost 2 km close to the sea and as an unconformity at depths of more than 1 km offshore. 8686 5.8 tectonic evolution of the west greenland epcm synthesis of information derived from sla, palaeothermal methods and the stratigraphic record of the west greenland margin has defined the tectonic development of its modern day continental margin (figs 38, 54). the present-day high topography was formed predominantly since the late miocene, by differential uplift (involving disruption into tilted blocks) and dissection of a peneplain, the remnants of which can now be recognised as plateaux at elevations up to almost 2 km close to the sea and as an unconformity at depths of more than 1 km offshore (fig. 55). our results show that an interval of 20 to 25 million years was sufficient to form a regionally extensive peneplain such as the ups, and imply that erosion over such a period of time was sufficient to erase older surfaces, where they were not protected by cover rocks. on the other hand, an interval of about 6 million years was insufficient time for the valleys of the lps to coalesce into a regional peneplain. the development of the west greenland margin involved multiple cycles of subsidence and burial and subsequent uplift and erosion (fig. 54). the formation of the elevated terrain in west greenland postdates the cessation of rifting by c. 50 million years, and the initial onset of uplift and erosion which led to formation of the ups close to sea level appears to correlate with the termination of sea-floor spreading in the adjacent ocean basin (srivastava 1978). these conclusions contradict the common assumption that continuous uplift, denudation and cooling histories are appropriate for epcms, and also suggest that the present landscape of this epcm is dominated by processes subsequent to rifting, break-up and continental separation. the similarity of the west greenland epcm to those in other parts of the world leads us in turn to speculate that the same might be true of other margins. we discuss evidence from a number of margins in support of this speculation in the next chapter. 87 6. contrasting views on the development of epcms in other areas from landscape studies and thermochronology 6.1 introduction the previous chapter demonstrated how integration of low-temperature thermochronology and landscape analysis with basic geological constraints defines the evolution of the continental margin of west greenland involving multiple episodes of subsidence/burial and subsequent uplift and denudation. in contrast, many published studies of other epcms which attempt to integrate results from these three disciplines within a common framework have been less successful. such studies have commonly been carried out within a paradigm of monotonic cooling and continual denudation, and have in some cases resulted in major inconsistencies between the different approaches. in this chapter we review evidence from epcms in southern africa, south-east australia and brazil, classic areas where extensive investigations have been carried out from both thermochronology and landscape analysis, and then briefly review results from other epcms. inconsistencies between interpretations derived from the two approaches in previous studies are highlighted, and the underlying reasons are investigated. evidence is presented which suggests that the evolution of many epcms was in fact more similar to that described for west greenland than previously envisaged, raising the possibility that this type of development may be the rule, rather than the exception. this has major implications for the nature of the underlying processes, as discussed in chapter 7. 6.2 southern africa 6.2.1 classic landscape studies southern africa is a classical area for landform analysis, with detailed studies resulting in the description of stepped low-relief surfaces graded to the general base level (sea level) formed at a successively upwarped margin (e.g. king 1967, 1972). in these works, and in more recent work by partridge & maud (1987), the emphasis is on a history involving continuous uplift and denudation since the eruption of the lower jurassic basalts of the drakensberg group (the youngest preserved stratigraphic unit of the upper carboniferous to lower jurassic karoo supergroup; e.g. tankard et al. 2009), which forms the highest topography in south africa (fig. 2). the focus in these studies is on definition and dating of regional planation surfaces representing the staged uplift of southern africa. king (1967) identified two surfaces above the drakensberg escarpment, namely gondwana and post-gondwana, and regarded them to be of jurassic and early-mid cretaceous age, while he thought the major summit planation surface below the escarpment and inland of it, the african surface, was formed from the late cretaceous to the early miocene with most erosion in the late cretaceous (table 3; cf. fig. 2). three lower landscape generations were interpreted to be the results of miocene and pliocene uplift episodes, by correlation with coastal or offshore deposits. partridge & maud (1987) came to a different conclusion to that of king (1967), finding no evidence for the existence of preserved mesozoic surfaces above the african surface (table 3). they regarded the african surface as being present on both sides of the escarpment, and to be the result of a single cycle of erosion from the time of rifting to the early miocene, with most erosion during the jurassic to cretaceous. they further identified two post-african surfaces and two neogene uplift events. in contrast to these earlier studies in which events were seen as extending continent-wide, moore et al. (2009) described the high-level plateau of southern africa as defined by three separate divides representing ‘axes of flexure‘ dating from early cretaceous, mid-cretaceous and palaeogene times. one aspect of the south african landscape that dominates much of the debate regarding the development of the present-day landscape is the great escarpment (e.g. king 1962). ollier & marker (1985, p. 49) succinctly described the major geomorphic features of southern africa as “a plateau... bounded by the great escarpment, and younger erosional features between the escarpment and the sea”. as discussed in section 2.2, ollier & marker (1985) interpreted the plateau as an old land surface (the 8888 ‘palaeo-plain‘), with significant denudation restricted to the region between the escarpment and the coast. this concept was developed further by ollier & pain (1997, p. 1) who described the palaeoplain as “little changed from the land surface that existed before continental break-up”, correlating with a ‘basal unconformity‘ in the offshore sedimentary section. within this concept, this elevated ancient surface has never been covered by a younger sedimentary section and has not undergone significant denudation. kempf (2010) provided an extensive review of alternative previous viewpoints in regard to the great escarpment of southern africa, pointing out that in addition to the view of the great escarpment as a remnant of rifting, others have regarded it as a post-rift feature. kempf (2010) concluded on the basis of morphological and geological evidence that the great escarpment of north and south namibia is the end-product of denudation resulting from tectonic uplift of the continent following rifting and break-up and formation of the south atlantic ocean. in these and other classical studies of southern african landscapes, amounts of denudation are limited essentially to those required for ‘infilling’ of the landscape to the level of these planation surfaces, and no consideration is given to the possibility that extensive sedimentary sequences may have been deposited and subsequently removed, in some cases even when sedimentary outliers are preserved. however, application of low-temperature thermochronology to southern africa provides a quite different view, as reviewed below. 6.2.2 low-temperature thermochronology studies numerous apatite fission-track studies across southern africa over the last 20 years (brown et al. 1990; gallagher & brown 1999a, b; brown et al. 2000, 2002; raab et al. 2002, 2005; kounov et al. 2008, 2009; tinker et al. 2008a) have established that the region has undergone major denudation during the cretaceous. a map of apatite fission-track ages, based on these studies, is shown in fig. 56. a more extensive dataset, showing similar features, is illustrated by gallagher et al. (1998), but full details of this dataset have not yet been published. across much of the region, including areas extending more than 500 km inland from the west and south coasts, apatite fission-track ages are less than 145 ma. these fission-track ages cannot be interpreted directly as indicating the timing of any specific event (section 4.1.6), but because the analysed samples are jurassic or older, these results can be regarded as indicating cretaceous (or later) cooling from palaeotemperatures around 100°c or above over a wide region. the cooling revealed by the apatite fission-track ages shown in fig. 56 has been explained by the authors of table 3. development of african palaeoplains as described by different authors king (1972, 1982) partridge & maud (1987) gondwana, jurassic uplift post-gondwana early cretaceous; proto-drakensberg uplift african surface, late cretaceous (main erosion) – early miocene; below drakensberg escarpment uplift 20 ma rolling landscape miocene sediments uplift; end of miocene widespread landscape, pliocene uplift youngest landscape quaternary (valley of a thousand hills) non-existing non-existing african surface, one cycle of erosion: rifting – early miocene; deep weathering profiles; most erosion: jurassic–cretaceous uplift post-african i miocene – end of pliocene uplift; end of pliocene post-african ii partridge (1998) non-existing non-existing massive denudation (2–3 km) during the early cretaceous resulting in the african surface; silcretes; eocene peripheral cover uplift, early miocene post-african i major neogene uplift post-african ii and gorges 89 the respective studies solely in terms of denudation following early cretaceous rifting, on the basis that denudation is the only mechanism capable of producing significant cooling at rifted margins (gallagher & brown 1997; 1999a, b; brown et al. 2000, 2002). in namibia, estimates of the total amount of section removed following rifting are around 3 km or more across a region extending 300 km or more inland and as much as 5 km at the coast (gallagher & brown 1999a, b; brown et al. 2000, raab et al. 2002, 2005), with a major phase of accelerated cooling in the late cretaceous (c. 70 ma). on the natal coast in the south-east (see fig. 2), brown et al. (2002) estimated that a minimum of 4.5 km of section has been removed since 130 ma. farther inland, but still seaward of the drakensberg escarpment, brown et al. (2002) estimated around 3 km of denudation since c. 91 ma, while west of the lesotho highlands, they estimated around 1.7 km of denudation since 78 ma, with accelerated denudation through the late cretaceous and much lower rates through the cenozoic at both locations. tinker et al. (2008a) interpreted their results in terms of two dominant phases of denudation, in the intervals 140 to 120 ma and 100 to 80 ma, with between 2.5 and 3.5 km of section removed in the late cretaceous phase. 0 250 500 750 1000 1250 1500 1750 2000 2250 2500 2750 3000 3250 elevation (m a.s.l.) 25°e 30°e15°e 20°e <60 aft ages (ma) 250 km fig. 58 60–90 91–120 121–180 181–250 >251 20°s 25°s 30°s africa fig. 56. map of elevation and published apatite fission-track ages across southern africa (see section 6.2.2 for sources). ages as young as 120 ma or less, which clearly post-date rifting and continental separation, are found hundreds of kilometres inland and bear no obvious relationship to the topography, including the location of the great escarpment. 9090 6.2.3 attempts to reconcile landscape studies and thermochronology in southern africa the results of these and many other low-temperature thermochronology studies of epcms are commonly discussed within a framework in which the modern topography is assumed to be related to continental rifting and break-up, and the onshore margins are assumed to have remained elevated since break-up, with significant denudation restricted to the region seaward of the great escarpment. gallagher et al. (1998) outlined three models embodying these concepts, as illustrated in fig. 57a–c. these models are expected to produce different patterns of apatite fission-track age variation (fig. 57d), which could be diagnostic in understanding the nature of the processes controlling the development of epcm landscapes. although these models have been widely employed as a conceptual framework for understanding low-temperature thermochronology data from passive margins (e.g. bishop 2007; campanile et al. 2008; burke & gunnell 2008), the results shown in fig. 56 do not display the patterns of age variation expected from these models. gallagher & brown (1999a, p. 48) concluded that “data are broadly consistent with a model where the majority of pre-break-up surface sea pre-rift rock syn-rift sediment post-rift sediment age of rifting original age distance from margin fi ss io ntr ac k ag e removed section a. downwarped rift shoulder b. elevated rift shoulder c. elevated rift shoulder with pinned divide d c b 0 0 0 a fig. 57. geomorphological models that are commonly used in discussions of the development of elevated, passive continental margins (epcms; after gallagher et al.1998). a: down-warped rift shoulder (ollier & pain 1997). b: scarp retreat from an elevated rift shoulder (e.g. gilchrist & summerfield 1991). c: down-wearing of an elevated rift shoulder with pinned divide (e.g. brown et al. 2002; persano et al. 2006). d: variation of fission-track age vs. distance expected from models a, b and c (red curves) compared with ‘original age’ (green) and ‘age of rifting’ (blue). in these models significant post-rift denudation (and hence apatite fission-track ages close to or younger than the time of rifting) is limited to the coastal side of the escarpment. but as shown in figs 56, 58 and 60, results from two much-studied epcms suggest that young fission-track ages, denoting significant post-rift cooling, extend much farther inland than the escarpment, and these models do not provide an accurate description of the denudation histories of epcms. facing page: fig. 58. variation of apatite fission-track age with elevation across the south-west african elevated, passive continental margin (epcm). a: map showing a close-up of the fission-track age map of southern africa in fig. 56. profiles b and c: data points (red dots) extending up to 500 km or more into the continent allow a direct comparison with the concepts represented in fig. 57 (data from tinker et al. 2008a and kounov et al. 2009). these panels are colour coded to match the representation of parameters in fig. 57d, viz: ‘original age’ (green), and ‘age of rifting’ (blue). three options are illustrated for the ‘original age’ which effectively corresponds to the time at which track retention began in the apatites. the most likely candidate is the extensive karoo igneous event (candidate 3 in panels b and c; 184–179 ma, duncan et al. 1997; svensen et al. 2012) while other possibilities are the cape orogeny (candidate 2), dated to either 278–230 ma (newton et al. 2009) or 250–215 ma (tankard et al. 2009) or the hiatus between deposition of the cape and karoo groups which tankard et al. (2009) assign to the interval 330–302 ma (candidate 1). ages of rifting on the southern and western margins are also indicated (candidates 4 and 5 from macdonald et al. 2003). measured fission track ages do not show the type of behaviour shown by any of the notional trends in fig. 57, and we conclude that these models do not provide an accurate depiction of the evolution of epcms. rsa: republic of south africa. 91 east-west north-south original age candidates (cf. fig. 57): 1: cape–karoo hiatus 330–302 ma 2: cape orogeny 278–230 ma or 250–215 ma 3: karoo terminal igneous activity 184–179 ma age of rifting candidates (cf. fig. 57): 4: opening of south atlantic 135–120 ma 5: rift basin development on southern margin 165–135 ma 0 50 100 150 200 250 300 2.0 0100500 200400 300 distance from coast (km) fi ss io ntr ac k ag e (m a) el ev at io n (k m a .s. l.) 1.5 1.0 0.5 0 2.0 el ev at io n (k m a .s. l.) 1.5 1.0 0.5 0 n am ib ia rs a great escarpment 60–90 samples from borehole apatite fission-track age and 2σ uncertainty 91–120 121–150 151–200 fission-track age (ma) tinker et al. (2008a) transect profile b profile c mossel bay cape town 20°e15°e 30°s 25°sa b c late jurassic – early cretaceous oudtshoorn basin town 1 2 3 4 5 1 2 3 4 5 0100500 200400 300 distance from coast (km) 0 50 100 150 200 250 300 fi ss io ntr ac k ag e (m a) oudtshoorn basin oudtshoorn basin 9292 denudation occurs in what is now the low-elevation coastal plain, seaward of the topographic escarpment…, with regional anomalies in the interior and along-strike of the margin related to post-breakup reactivation of regional structures.” but as is clear from fig. 56, ages as young as 100 ma or less occur up to 500 km inland, well away from any structures that may have been reactivated. this is further emphasised in fig. 58, where measured fission-track ages from the south-west of south africa along two transects shown in fig. 56 are plotted against distance from the coast, together with the respective topographic profile along the transect. for purposes of comparison with fig. 57, various candidates for the ‘original age’ are shown in green, and two possibilities for ‘age of rifting’ are shown in blue, while data are shown in red to match the colour coding of lines in fig. 57. measured fission-track ages are consistently lower than the various candidate ‘original ages’ across the whole length of both transects, over distances of over 500 km. the conceptual models illustrated in fig. 57 therefore provide little guide to the processes actually responsible for the observed pattern of fission-track ages across southern africa. estimates of the amount of section removed by cretaceous denudation derived from apatite fissiontrack data have traditionally proven difficult to reconcile with results from studies based on analysis of landforms. swart (2006) has drawn attention to these problems in namibia, considering interpretations of apatite fissiontrack data in terms of up to 5 km of post-early cretaceous denudation at coastal locations and up to 3 km or more inland to be totally unrealistic on the basis of the preserved geological section. aizawa et al. (2000) provided a view of the evolution of onshore namibia based on stratigraphic and geomorphic constraints which bears little relation to the history suggested by the interpretations derived from the apatite fission-track data, with only small thicknesses of section removed since continental break-up. some syntheses of onshore denudation histories based on such ‘conventional approaches’ (bluck et al. 2005; goudie 2005) fail to mention the apatite fissiontrack data and their interpretations at all, highlighting the disparity between conclusions derived from the two approaches and the difficulties of reconciliation. fission-track ages as young as 100 ma or less in samples from inland of the great escarpment in fig. 56 are incompatible with the idea that the inland plateau represents a pre-rift land surface (ollier & pain 1997) or any trace thereof. partridge (1998) attempted to reconcile the requirement for 3 km of missing section to explain apatite fission-track data from south africa with geological and geomorphological information by suggesting that all erosion was essentially completed by the end of the early cretaceous. however, as highlighted above, this is inconsistent with the timing of denudation defined by these apatite fission-track studies, which require significant amounts of late cretaceous denudation. studies of the morphology of kimberlite pipes also suggest much lower amounts of denudation than indicated by the apatite fission-track interpretations, suggesting a maximum of c. 1.8 km of post-early cretaceous denudation (hawthorne 1975), while studies of different types of xenoliths in kimberlites of different age suggest less than 1 km of denudation since c. 85 ma over most of the region (hanson et al. 2009). burke & gunnell (2008) proposed a two-stage model to explain the pattern of apatite fission-track ages across southern africa, involving scarp retreat from initial rifting (similar to fig. 57a) followed by uplift of a “great swell” and resulting erosion (equivalent to fig. 57b) over the last 30 million years. burke & gunnell (2008) suggest that this will produce a pattern similar to trend c in fig. 57d, referencing gallagher et al. (1998) as suggesting that this trend provides the best description of the measured age variation. the rationale behind this two-stage model is not clear, as burke & gunnell (2008) also say that erosion over the last 30 million years in their scenario will not have much impact on the apatite fission-track data. so it remains unclear how a pattern of ages similar to trend c in fig. 57d, which shows partially reset ages extending inland of the great escarpment, will result from this model. as shown in fig. 56 neither the model espoused by burke & gunnell (2008) nor any of the models illustrated in fig. 57 can explain the observed pattern of apatite fission-track ages, which shows pervasive, uniform, young fission-track ages close to or younger than the time of break-up extending hundreds of kilometres inland from both the southern and western continental margins. 6.2.4 possible alternative explanations in considering why apatite fission-track studies and other approaches lead to such disparate views of the denudation history of south-west africa, one factor that requires reassessment is the assumption in previous studies that denudation is the only process that causes significant cooling at continental margins (i.e. palaeotemperatures solely represent greater depth of burial, and palaeogeo93 thermal gradients were similar to present-day values). elevated heat flow associated with rifting would significantly reduce the amount of additional burial (and subsequent denudation) required to explain the palaeotemperatures indicated by the apatite fission-track data. afta data from a number of margins, including southeast australia (duddy et al. 1994; duddy 1997; green et al. 2004), west greenland (japsen et al. 2005, 2006) and west africa (bray et al. 2002; turner et al. 2008), have provided abundant evidence of significantly elevated palaeogeothermal gradients, in most cases associated with continental rifting and break-up. however, this is far from universal, as noted at the uk north atlantic margin by green et al. (1999) and the atlantic margin of brazil by japsen et al. (2012b; section 6.3). evidence in support of elevated cretaceous heat flow along the atlantic margin of south-west africa was reported by whitehead et al. (2002) who concluded, from the mineralogy of mantle xenoliths in upper cretaceous intrusives, that the late cretaceous heat flow in coastal locations was twice the value in the vicinity of the gibeon kimberlites 300 km inland. in addition, tinker et al. (2008b) reported that while the timing of denudation inferred from apatite fission-track data in outcrop samples from the south-western part of south africa (fig. 58) showed a good match to that of accelerated deposition offshore, the amount of sediment preserved in offshore basins is less than half of that expected on the basis of the denudation required to explain the apatite fissiontrack data. this mismatch can be readily explained if the cretaceous heat flow was higher than the present day value, although erosional removal of offshore sediments (e.g. mcmillan 2003) also contributes to some degree in explaining the discrepancy. tinker et al. (2008b) also drew attention to the temporal coincidence between episodes of cretaceous igneous activity and enhanced denudation in south africa. the contribution from intrusive activity to the thermal histories revealed by apatite fission-track data in namibia has not been considered in published studies to date. numerous episodes of igneous activity are recognised across namibia, including lower cretaceous anorogenic complexes (e.g. milner et al. 1995), upper cretaceous bodies such as the gross brukkaros structure and the regionally extensive gibeon kimberlite field with ages of c. 77 ma and between 72 and 79 ma, respectively (reid et al. 1990), as well as a number of alkaline plugs (whitehead et al. 2002). cenozoic igneous activity includes the klinghardt phonolites around luderitz (c. 37 ma, marsh 1975; lock & marsh 1981). possible heating mechanisms that may be manifested in the afta data include elevated heat flow, local exhumation related to thermal doming or hydrothermal effects, which are particularly pronounced around the gross brukkaros intrusion (miller 2008). afta and vr data from the heavily intruded sequences of the uk atlantic margin have shown that the thermal effect of intrusive bodies can be more widespread than that expected on the basis of simple conductive heating (parnell et al. 1999; duddy et al. 1994, 1998). by comparison it seems likely that the multiple episodes of igneous activity that have affected namibian coastal regions may well have produced profound palaeothermal effects, which may have been interpreted in terms of deeper burial in previous studies. a key factor in many of the published apatite fission track studies of southern africa is that they do not take into account geological constraints on the underlying thermal and denudational histories, and in many cases the interpretations derived from apatite fission-track data appear to conflict with geological evidence. for example, late cretaceous cooling reported by raab et al. (2002) in outcrop samples from northwest namibia in and around the damara fold belt was interpreted solely in terms of denudation (as above), involving removal of several kilometres of section. such an explanation is difficult to reconcile with geological evidence, given the presence in that area of outcropping jurassic sediments at waterberg and mt etjo (see miller 2008), as well as the lower cretaceous etendeka volcanics closer to the coast, which show that the present-day surface was also close to the surface during the mesozoic. this implies that much of the rock that was removed during late cretaceous exhumation across the region must have been first deposited on top of these mesozoic units before being subsequently removed. recently, dauteuil et al. (2013) presented a history of the namibia margin in which the palaeothermal effects defined from the apatite fission-track data were explained by deeper burial in the early cretaceous by up to 5 km of etendeka volcanics, subsequently eroded in the late cretaceous. this explanation is more compatible with stratigraphic constraints than removal of basement or karoo supergroup rocks as suggested by raab et al. (2005), although it appears to be open to speculation as to whether the etendeka basalts ever attained such thicknesses. inconsistencies between the results of apatite fission track and landscape studies in southern africa should not detract from the evidence in fig. 56 that rocks now 9494 outcropping across much of southern africa were at 100°c or more during the cretaceous prior to the onset of exhumation. since such palaeotemperatures are unlikely to occur within 1 km of the surface for any reasonable geothermal gradient, these data suggest that significant amounts of section must have been removed across the region, and geomorphological studies that fail to account for the cooling reflected in the apatite fission-track data cannot be considered realistic. this implies that the major planation surfaces that now define the elevated topography across southern africa were formed long after rifting and break-up, by erosion that included removal of substantial thicknesses of sedimentary cover. plausible interpretations of the apatite fission-track data across southern africa will only emerge by considering the thermal history data within the context imposed by independent geological and geomorphological constraints. as we illustrate below, consideration of existing data in this light suggests a more complex evolution than the continuous denudation histories employed in all the studies discussed above. 6.2.5 evidence for post-breakup subsidence and burial of the southern africa margin all of the apatite fission-track and geomorphological studies referred to above are interpreted in terms of continual cooling and long-term denudation. however, as illustrated here, data presented by tinker et al. (2008a) from south africa suggest that the development of the southern margin of africa may have been more similar to that of west greenland presented in chapter 5. tinker et al. (2008a) reported apatite fission-track data in outcrop samples, collected along a roughly northsouth transect (fig. 56), extending from the south coast and crossing the great escarpment into the elevated inland region to a point c. 500 km from the coast. the region south of the escarpment contains considerable relief, making up the mountains of the cambrian to carboniferous cape supergroup reaching over 2000 m a.s.l., but the erosional base of the landscape descends from about 800 m a.s.l. south of the escarpment to sea level at the coast. tinker et al. (2008a) analysed samples of the cape supergroup and older (pre-cape) units in the south, and various permian and triassic units of the karoo supergroup in the north of their transect, plus additional samples from three deep boreholes in the karoo sequence south of the escarpment. results from these samples were interpreted as defining major late cretaceous (between 100 and 80 ma) cooling, while samples from a fourth borehole north of the escarpment were interpreted as showing an early cretaceous phase of cooling. tinker et al. (2008a) attributed late cretaceous cooling to denudational removal of a once much thicker karoo sequence extending in time to the jurassic drakensberg volcanics (c. 183 ma; duncan et al. 1997). for karoo supergroup samples from the northern part of their transect (fig. 58c), tinker et al. (2008a) showed that the required thickness of former cover can easily be accommodated by the thicknesses of younger karoo units preserved elsewhere in the basin. but at locations in the south of the transect, the presence of upper jurassic to lower cretaceous sedimentary units of the uitenhage group, deposited mainly between c. 151 and c. 135 ma (shone 2009), in extensional basins such as the oudtshoorn basin (location shown in fig. 58), rules out such an interpretation. while the northern margins of these basins tend to be fault-controlled, at southern basin margins and elsewhere the sedimentary units of the uitenhage group rest directly on palaeozoic metasedimentary rocks of the cape supergroup and/or older basement. therefore, even if any units of the karoo basin sequence were deposited in this region, for which there is no evidence (johnson et al. 2009), they must have been removed prior to the late jurassic. late cretaceous (100 to 80 ma) temperatures of up to c. 100°c are expressed in the fission-track data of tinker et al. (2008a), in outcropping samples of cape supergroup and older units across this region, including one sample close to the southern margin of the oudtshoorn basin where uitenhage group sediments are in depositional contact with underlying units. these rocks must have been at or close to the surface in the late jurassic to early cretaceous, and therefore burial of the palaeozoic and older rocks by a much thicker sedimentary cover post-dating the preserved sedimentary units of the uitenhage group is required in order to achieve the estimated temperatures prior to the onset of late cretaceous denudation. subsequent work (green et al. 2011) has shown that the outcropping uitenhage group sedimentary rocks across the south-western part of south africa were also heated to palaeotemperatures around 80°c or more prior to exhumation which began between 85 and 70 ma. these results confirm that the region was buried by up to 2 km of lower cretaceous section prior to late 95 cretaceous exhumation. the presence of sedimentary remnants such as these late jurassic-early cretaceous extensional basins along the south coast of south africa therefore provides a key geological constraint on the likely denudation history of the region. none of the published apatite fissiontrack studies of southern africa have taken account of such constraints, and as a result these studies have failed to reveal the true nature of the development of the margin. it is also worth pointing out that as illustrated in fig. 59, the interpretation by tinker et al. (2008a) that a thickness of c. 3 km or more of section was removed since the late cretaceous from the region below the great escarpment, where the boreholes are located, implies that unless massive faulting or flexure can be invoked, the great escarpment must also have been buried by around 2 km of section at that time. and if an early cretaceous phase of denudation affected this region as suggested by tinker et al. (2008a), the present-day summit of the great escarpment must have been buried even more deeply at that time. this makes it difficult to envisage a scenario in which the great escarpment represents a fundamental control on post-rift denudation as embodied in the comparison of eroded thicknesses with offshore sedimentation by tinker et al. (2008b). instead, these observations suggest that the present-day great escarpment is the end result of the denudational history, rather than representing a fundamental control on the process of denudation. its location today appears to owe much to the presence of resistant karoo dolerite sills. 6.2.6 summary integrating geological constraints with published apatite fission-track data shows that the development of the modern topography characterising the southern margin of africa was much closer to that of west greenland outlined in chapter 5 than previously envisaged. following break-up, the region subsided and was buried by up to 2 km of section, which was removed by uplift and denudation that began much later. the present-day topography is therefore not related to processes of continental rifting or break-up, but is instead the result of post-breakup processes, similar to the conclusion of kempf (2010) in relation to the escarpment of namibia (section 6.2.1). 6.3 south-east australia 6.3.1 thermochronology the classic study by moore et al. (1982) of apatite fissiontrack data from south-east australia was one of the earliest applications of low-temperature thermochronology to a rifted margin, defining for the first time a now-familiar pattern of young apatite fission-track ages (around 100 ma) along the coast, increasing inland to values around 300 ma in the interior highlands. similar patterns of apatite fission-track ages have since been identified in many other areas of the world, as reviewed e.g. by van der beek et al. (1995), but as more data have become available it has become clear that in many areas young ages also occur hundreds of kilometres inland, as in figs 57, 60. kohn et al. (2002) compiled available apatite fissiontrack data from south-east australia (fig. 60) and, after rejecting the possibility that elevated heat flow played any significant role in the heating revealed by the data, interpreted their results as reflecting only heating due to additional depth of burial and cooling due solely to denudation, with up to 3 km of section removed from coastal locations since 130 ma. but even 100 km or more inland, kohn et al. (2002) suggest that around 1–2 km of section has been removed since c. 130 ma. this, together with apatite fission-track ages around 100 ma or fig. 59. conceptual representation of the northern portion of the n–s profile in fig. 58c, around the position of the great escarpment. tinker et al. (2008a) interpreted their results from boreholes south of the escarpment (shown in fig. 58) in terms of late cretaceous denudation, during which around 3 km of section was removed. this diagram emphasises that this interpretation requires that the present-day escarpment was buried by c. 2 km of section prior to the onset of late cretaceous exhumation. it is therefore clear that the present-day landscape is the product of the late cretaceous denudation history, and preserves no information on the pre-rift configuration (cf. fig. 57). 3 km3 km n s 1 2 3 el ev at io n (k m a .s. l.) present-day land surfacepresent-day land surface land surface prior to onset of late cretaceous denudation according to tinker et al. (2008a) great escarpment 9696 less many hundreds of kilometres inland in fig. 60, again emphasises the failure of the conceptual models in fig. 57 to describe real data. 6.3.2 landscape analysis the landscapes of south-east australia have been a subject of intense study over the last 100 years or more. in earlier morphological interpretations the high plains of 35°s 30°s 145°e250 km 150°e 0 250 500 750 1000 1250 1500 1750 2000 2250 2500 2750 3000 elevation (m a.s.l.) <60 aft ages (ma) 60–90 91–120 121–180 181–250 >251 australia fig. 60. map of published apatite fission-track ages across south-east australia (kohn et al. 2002). most of these results represent analyses of rocks of devonian age or older. as in southern africa, young ages are found hundreds of kilometres inland, and again bear no obvious relationship to the topography, including the location of the great escarpment. 97 south-east australia (regions of low relief at elevations above 1 kilometre) were regarded as evidence of late pliocene uplift following planation at sea level (andrews 1910) or alternatively oligo–miocene uplift (king 1967). craft (1933) also interpreted the incised valleys (corresponding to the great escarpment; ollier 1982) as witness of pliocene uplift. over time, opinions shifted towards regarding the high plains as old, remnant features, and interpretations favouring recent uplift were replaced by the idea that the south-eastern highlands represent the remnants of the palaeozoic lachlan fold belt (lambeck & stephenson 1986). more recently, opinion has once again shifted towards emphasising cenozoic uplift (holdgate et al. 2008), assisted by a growing body of evidence of late cenozoic tectonism across south-east australia (dickinson et al. 2002; sandiford 2003). bishop & goldrick (2000) provide an extensive review of previous work, and comment on the particular interest of the eastern australian margin in view of its ‘general tectonic stability’ and ‘generally low rates of denudation’. a key aspect of geomorphological investigation in south-east australia has been the quest to identify landscape elements of mesozoic age (as reviewed by bishop & goldrick 2000; also see for example hills 1975; hill 1999), coupled with the interpretation of aspects of the present-day landscape as representing long-term stable axes (persano et al. 2006). another strong emphasis has been on recognition of the impact of rifting on the landscape, with the great escarpment (high margin with deeply incised valleys; fig. 3) commonly interpreted as a direct consequence of the rifting process, as also thought by many to be the case in south africa (e.g. ollier 1982, 1985, 1995, ollier & marker 1985; ollier & pain 1997). 6.3.3 attempts to reconcile thermo chronology and landscape analysis the notion of south-east australia as tectonically stable does not fit comfortably with the interpretation of kilometre-scale denudation suggested by apatite fissiontrack data (above). as summarised succinctly by bishop & goldrick (2000, p. 246): “the various thermochronologic approaches ... suggest that the present land surface is ... the result of kilometre-scale denudation in the late mesozoic, whereas some geomorphological interpretations indicate the preservation of mesozoic landscape elements at the present land surface”. on this basis, it is very difficult to reconcile these conflicting views whilst remaining faithful to both. the apatite fission-track data summarised in fig. 60 again provide no support for the conceptual models shown in fig. 57, with ages as young as 120 ma or less in palaeozoic basement rocks hundreds of kilometres inland from the coast. in addition, apatite fission-track data (o’sullivan et al. 1995) and vr and palaeomagnetic data (middleton & schmidt 1982), show that the early cretaceous palaeothermal signature along the southeastern australian margin shows no obvious correlation with the presence and location of the coastal escarpment. the interpretation by middleton & schmidt (1982) that their vr and magnetisation measurements indicated up to 3 km of additional section at coastal locations in the permian–triassic sydney basin reawakened an old debate. branagan (1983) reviewed the history of debate regarding the ‘vanished sequence’ of the sydney basin, and concluded on the basis of evidence from preserved stratigraphy and geomorphology that less than 1 km of section could have been deposited and subsequently removed in post-triassic time. more recently, apatite fission-track data from southeast australia have been the focus of further discussion and controversy (e.g. nott & purvis 1995; leaman 2003; osborne et al. 2006) without any satisfactory resolution being reached. geomorphological studies tend to focus on continuous denudation, and interpret the absence of any sedimentary cover as indicating that none was ever deposited. on the other hand, thermochronological studies have historically tended to ignore constraints imposed by geology. a recent review from a geomorphological perspective of methods for dating land surfaces, drawing heavily on australian examples (watchman & twidale 2002) fails to mention apatite fission-track methods, despite including a wide range of other methods from radiocarbon to ar-ar dating, perhaps illustrating a scepticism surrounding apatite fission-track-based interpretations in general. as with south-west africa, one aspect of the apatite fission-track interpretation that may help to bring these opposing views closer together is the possibility of significantly elevated heat flow associated with rifting along the south-east australian margin, which would reduce the amount of removed section required to explain the apatite fission-track data. middleton & schmidt (1982) comment that vr data require higher palaeogeothermal gradients at coastal locations in sydney basin wells, and elevated early cretaceous palaeogeothermal gradients have been documented in the otway basin (duddy 9898 1994, 1997; cooper 1995; cooper & hill 1997; green et al. 2004), so it seems reasonable to suppose that this may also have been the case along the south-east australian margin. nevertheless, as with the case of south-west africa, for any reasonable value of palaeogeothermal gradient, kilometre-scale denudation is required to explain the apatite fission-track data, both along the coastal strip but also in the highland regions. to date, no convincing mechanisms have been proposed by which this can be achieved, and this has no doubt contributed to the difficulty in accepting this interpretation. persano et al. (2002, 2005) attempted to integrate thermochronological data more directly with evidence from geomorphology in south-east australia. persano et al. (2002) suggested, based on new apatite (u-th)/he ages from the same region of south-east australia originally studied by moore et al. (1982), that the upland plateau region had undergone only limited denudation (<1 km of removed section) over the last 180 million years, compared to much greater denudation of the coastal plain region (several kilometres of removed section) since break-up in the early to mid-cretaceous. this contrasts markedly with previous interpretations derived from apatite fission-track data (above) that the upland region has also undergone kilometre-scale denudation since break up. but as reviewed earlier (section 4.2), it has recently become clear that interpretation of apatite (u-th)/he ages is not as straightforward as previously assumed. thus, it is likely that the reason for this mismatch between the interpretation from apatite fission track and the conclusions derived by persano et al. (2002, 2005) from their apatite (u-th)/he ages probably arises because their apatites have retained more he than expected due to increased levels of radiation damage. hence, in terms of the systematics used in their interpretation the apatite (u-th)/he ages are ‘anomalously old’. we should note that persano et al. (2005) did consider the possibility of elevated heat flow along the margin, related to rifting, in contrast to previous studies. in summary, despite large amounts of apatite fissiontrack data from south-east australia and over 100 years of geomorphological investigation, no reconciliation has yet been achieved between the two approaches or within them. the nature of the processes responsible for the heating and cooling revealed by the apatite fission-track data remains the subject of debate, which has no doubt contributed to the lack of reconciliation with other lines of evidence. firm constraints on palaeogeothermal gradients at the early cretaceous palaeothermal maximum, together with a precise chronology for the onset of cooling, would provide a major boost to the understanding of the underlying processes. to echo the comments of bishop & goldrick (2000), kilometre-scale denudation as required by apatite fission-track data from the upland region is very difficult to reconcile with the preservation of mesozoic landscape elements, and true reconciliation of the two approaches will require multi-disciplinary investigations in which unambiguous independent constraints are available to constrain interpretations from both approaches, as in the west greenland example discussed in chapter 5. quartz gravel mudstone basalt coal fig. 61. eocene muds overlying former river gravels and covered by eocene basalt near the summit of mount hotham in the south-east australian highlands. as discussed by holdgate et al. (2008), the presence of these muds at an altitude of c. 1800 m above the presentday sea level is strong evidence of post-eocene uplift of the present-day mountains. 99 6.3.4 evidence for episodic burial and uplift on the south-east australian margin interpretation of landscape elements as dating from the mesozoic in previous landscape studies of southeast australia is based chiefly on inference and unconstrained extrapolation of unconformity surfaces. recently, holdgate et al. (2008) drew renewed attention to the presence of eocene muds and sands preserved below a cover of basalts (former valley fills) forming summits of high-level plains at almost 2 km a.s.l. in the south-east australian highlands (fig. 61). holdgate et al. (2008) highlighted the lithological and other similarities between these eocene sediments and deposits of similar age in the subsurface within the gippsland and murray basins, to the south and north of the present-day mountains. they suggested that these similarities indicated that the eocene sediments were deposited within a lowrelief landscape at or near sea level, which was uplifted and dissected in post-eocene times. these observations suggest that the mountains of south-east australia are relatively young (post-eocene), favouring the earliest views on this subject, as reviewed in section 6.3.2. this scenario is strikingly similar to that outlined for west greenland in chapter 5. the firm constraints on the timing of uplift and dissection provided by holdgate et al. (2008) offer the promise of a more rigorous integration of information from thermochronology and landscape studies in the future. while not relating directly to post-break up events, reset fission-track ages and high vr values in sedimentary rocks of triassic age in the sydney basin (o’sullivan et al. 1995; middleton & schmidt 1982) emphasise that heating to the palaeotemperatures required to explain the apatite fission-track and vr data must have involved additional burial by upper triassic to lower cretaceous sediments which have been subsequently eroded away. the amount of additional section required is uncertain because of the lack of rigorous constraints on palaeogeothermal gradients. but the precise amount of additional burial is not as important as the fact that burial is required prior to the onset of cooling (i.e. denudation). thus, thermal histories involving monotonic cooling, as favoured by most thermochronological studies (as discussed earlier) are not relevant to this situation. this argument can be extended to coastal regions beyond the southern limit of the sydney basin sedimentary cover, where basement outcrops must also have been buried by a permian to lower cretaceous sedimentary cover. this emphasises the importance of sedimentary outliers in helping to define the evolution of the margin, as illustrated elsewhere throughout this paper. 6.4 brazil 6.4.1 introduction the atlantic margin of brazil is a classic area for studies focused on the origin of the extensive plateaux that characterise the country, even far from the coast (fig. 62; e.g. king 1956a, 1967). plateaux – or planaltos (elevated plain) in portuguese – are key aspects of the brazilian geography: the capital brasilia was developed on such a planalto about 1200 m a.s.l., and today the address of the web site of the brazilian government is ‘planalto.gov.br’. the geological record in north-east brazil allows direct evaluation of the vertical motion of the margin following break-up of the south atlantic at the aptian–albian transition (torsvik et al. 2009). the highly fossiliferous santana formation – in particular famous for fossil fish – crops out on the flanks of the araripe plateau. it is of early albian age and part of the post-rift succession of the extensional araripe rift basin (fig. 63; assine 2007). the santana formation is represented largely by nonmarine or quasi-marine strata, but according to martill (2007), the occurrence of echinoids (sea urchins) in the santana formation, defining an unambiguously marine horizon, has been known and documented since 1966. arai (2000) also reported additional indicators of midcretaceous marine environment across the interior of brazil. the presence of these marine, post-rift strata in the interior of brazil at elevations up to 600 m a.s.l. (morais neto et al. 2006) testifies to both post-rift subsidence and subsequent significant uplift. furthermore, morais neto et al. (2006) presented afta data which document that post-rift sediment from the araripe plateau reached palaeotemperatures of 80–100°c in the late cretaceous, and thus that burial of the araripe rift continued after the deposition of the cenomanian deposits that are the youngest sediment preserved today (assine 2007). despite this clear evidence of post-rift subsidence and much later uplift and erosion, scientific controversy abounds in the literature surrounding the antiquity of landscapes on the brazilian epcm. many brazilian as well as overseas authors emphasise that the formation of the present-day landscape took place millions of years after break-up (e.g. king 1956a, 1967; almeida http://planalto.gov.br 100100 and carneiro 1998 (and references therein); valadão 1998; ab’saber 2000; cobbold et al. 2001; riccomini et al. 2004; zalán & oliveira 2005). other workers follow the most widely used approach at the present day, as described in section 2.4, in claiming that the dominant features in the landscape of brazil date back to rifting and break-up of the atlantic margin (e.g. gilchrist & summerfield 1990; gallagher et al. 1994, 1995, 1998; ollier & pain 1997; sacek et al. 2012). we seek to resolve the apparent conflict between these opposing views in the following discussion. 6.4.2 studies focused on landscape evidence lester king (king 1956a, 1967) presented the most complete geomorphological analysis of eastern brazil so far. king identified surfaces from a combination of fieldwork and early topographical maps and used geological evidence to place time constraints on the palaeosurfaces he mapped (e.g. king 1967, p. 233). he regarded the highest surface in the landscape as the oldest and the lower surfaces to be younger, and identified four denudation surfaces in eastern brazil, which he interpreted as representing cyclic, low-relief, base level-governed erosional surfaces. he interpreted the age of the two oldest (and highest) surfaces in eastern brazil to be of mesozoic age, and he named them gondwana and post-gondwana. he considered the gondwana surface to be a pre-break-up surface that had developed across south america and africa. in brazil, the gondwana and post-gondwana surfaces are only preserved in the highest areas in the interior; e.g. small remnants in chapada diamantina and in parts of planalto da conquista (see fig. 64). more extensive, according to king (1967), is the sul-americana (south america) surface that is well preserved, dominating the highlands of chapada diamantina as well as parts of the planalto da conquista. king attributed an early cenozoic age to the sul-americana surface because it cuts across silicreted sands of early cenozoic age in locations west of the são francisco river. king (1967) named a smooth surface with occasional inselbergs the velhas surface; e.g. on the flanks of planalto da conquista, which he regarded to be late cenozoic. king (1956a, 1967) assigned large parts of coastal and nearcoastal areas to belong to the youngest erosional cycle, the paraguaçu, and he thought that this cycle was due to recent tilting along the present coastline. peulvast & claudino-sales (2004) identified two erosional levels of regional extent in north-east brazil: a low plain between 0 and 300 m a.s.l. (the sertaneja surface or sertão) and the discontinuous remains of a high plain between 750 and 1100 m a.s.l.; including the araripe and borborema plateaux. peulvast et al. (2008) suggested that these plateaux are remnants of a continuous and low-lying, late cretaceous rift flank which was uplifted in postcenomanian time, following deposition of the youngest rocks now on the plateau. however, valadão (1998) regarded the stepped landscapes in north-east brazil to reflect mainly the cenozoic development. valadão (1998) thus found that the plateau surfaces in eastern brazil coincide with the regional erosion surface that truncates maastrichtian sediments in the sanfranciscana basin (fig. 62; campos & dardenne 1997). bonow et al. (2009) and japsen et al. (2012b) analysed landforms in north-east brazil (10–15°s; figs 64, 65) based on a digital elevation model from which they constructed a contour map (cf. bonow 2004; bonow et al. 2006a). the mapping was supported by profiles along rtj rift rio de janeiro brasilia rio de janeiro sf basin 10°s 20°s 250 km 40°w mesozoic rocks town 45°w fig. 64 araripe plateau brasilia salvador borborema plateau brazil elevation (m a.s.l.) 0 500 1000 1500 2000 fig. 62. elevation and extent of mesozoic rocks in eastern brazil. note the scattered occurrence of plateaux even far from the atlantic margin. sf: são francisco. rtj: recôncavo-tucano-jatobá. 101 a square grid where topographical profiles were plotted along the grid lines together with maximum and minimum heights within a swath. in this way they identified two surfaces of low, relative relief and of regional extent, the higher surface (hs) and the lower surface (ls); cf. figs 65, 66. there is also a coastal plain of limited extent, and above the level of the hs there are distinct hills along ridges of particularly resistant rock. the hs is preserved on high ground and includes the plateaux of chapada diamantina and planalto da conquista. these plateaux define a coherent surface, dipping slightly seawards. the hs cuts across precambrian rocks that are frequently deeply weathered. the hs and the cenozoic laterites (cprm 2003) are preserved where the drainage system is unaffected by the fluvial system related to the younger ls at lower elevations. the ls is well defined across wide areas in the interior. the ls cuts across precambrian basement in the west, the intracontinental recôncavotucano-jatobá (rtj) rift, and precambrian basement again in the east. it truncates the post-rift sediment of the aptian marizal formation within the rift (silva et al. 2007), but it also truncates the lower miocene, marine sabiá formation within the recôncavo basin (viana et al. 1971). bonow et al. (2009) found that both the hs and the ls are erosional features (peneplains) because they cut across rocks of different ages and resistance to erosion. the hs must originally have extended across a larger area than that of the present-day plateau remnants, as the ls has developed at its expense. the authors used geological constraints to conclude that the ages of the hs and ls are palaeogene and neogene, respectively, and that sea level is the most likely base level to which the surfaces were graded. the ls is rapidly being dissected by river incision into the sediment of the rtj rift due to a recent change in base level and thus producing a coastal plain. a palaeogene age for the hs across the highlands of chapada diamantina and planalto da conquista is consistent with the interpretation of king (1967) who mapped these plateaux as part of the palaeogene sulamericana surface. the highly dissected hs on the flanks of planalto da conquista corresponds to the neogene velhas surface of king, whereas the ls and the coastal plain are equivalent to his paraguaçu surface within the study area. further north, peulvast et al. (2008) also identified two erosional levels of regional extent: a surface at low elevation and the discontinuous remains of a high plain (including the araripe and fig. 63. profile across the araripe plateau (see fig. 62) with the early cretaceous rift section above jurassic and palaeozoic pre-rift sediments (assine 2007). the post-rift santana formation of early albian age contains unambiguous marine strata (e.g. echinoids) at an elevation of c. 600 m a.s.l. (morais neto et al. 2006; assine 2007; martill 2007). these observations testify (1) to post-rift subsidence and (2) to subsequent significant uplift. thermal history interpretation of afta samples from the post-rift sequence of the araripe basin shows that this sediment reached palaeotemperatures of 80–100°c in the late cretaceous in close agreement with vitrinite reflectance values of 0.56% corresponding to a palaeotemperature of 93°c (morais neto et al. 2006). these observations show that burial of the araripe rift continued after the deposition of the cenomanian sediments that are the youngest preserved today. the thickness of the cover removed above the present-day surface of the plateau since the onset of late cretaceous denudation thus amounts to c. 2 km (1.8–2.5 km for a palaeogeothermal gradient of 30°c and a palaeosurface temperature of 25°c). based on morais neto et al. (2006). post-rift sediment (aptian–cenomanian) afta sample c. 90°c in late cretaceous c re ta ce ou s co ve r 2 km afta sample c. 90°c in late cretaceous c re ta ce ou s co ve r 2 km n s -0.5 0 0 0.5 1 1.5 syn-rift sediment (valanginian–hauterivian) pre-rift sediment (jurassic and ordovician–silurian) precambrian basement 2 2.5 el ev at io n (k m a .s. l.) 5 km santana fm with marine sediment land surface prior to onset of late cretaceous denudation 102102 w e ? ? 0 1 2 0 1 2 precambrian basement jurassic–cretaceous sediment maximum altitude in corridor minimum altitude in corridor higher surface lower surface x x x' x' y y' y'y c 38°w39°w40°w 39°w40°w41°w42°w el ev at io n (k m a .s. l.) el ev at io n (k m a .s. l.) well location water divide n ot m ap pe d limit of rift basin photo location, fig. 66 a b c chapada diamantina planalto do maracás recôncavo basin sabiá fm (projected) b fig. 65 75 km cenozoic laterite barreiras fm (neogene) neogene sediment lower miocene sediment mesozoic sediment palaeozoic sediment precambrian rock 38°w40°w42°w 12°s 14°s 38°w40°w42°w 12°s 14°s a sabiá fm sabiá fm 75 km pm cd pc pm cd pc pm cd pc >300 fission-track age (ma) 250–300 75–130 180–250 130–180 well a well b fig. 64. a: geology and b: elevation of the area of north-east brazil studied by bonow et al. (2009) and japsen et al. (2012b). location shown in fig. 62. note that the presence of the early cretaceous rift only has a limited expression in the landscape, and that the plateaux coincide with extensive laterite covers. based on geological maps of brazil and of the state of bahia (cprm 2001, 2003). c: two profiles to illustrate mapping of low-relief surfaces and interpreted peneplains, i.e., higher surface and lower surface (bonow et al. 2009). plateaux are part of one, seaward-dipping surface (profile yy’). locations of profiles shown in b. mapped, low-relief surfaces cut across basement (consisting of rocks that have different resistance to erosion) as well as the sedimentary sequence of the recôncavo-tucano-jatobá rift. hence, surfaces are erosional features. the higher surface (hs) formed as a peneplain by fluvial erosion to near sea level during the eocene–oligocene. the hs was subsequently buried at the oligo–miocene transition and then uplifted and re-exposed in the miocene. uplift caused rivers to incise, after which the lower surface (ls) formed below the hs. cd: chapada diamantina. pc: planalto da conquista. pm: planalto de maracás. the location of the marine deposits of the lower miocene sabiá formation is indicated. modified from japsen et al. (2012b). 103 borborema plateaux). there are thus two regional peneplains in north-east brazil, a lower surface and remnants of a higher surface that were formed by erosion during neogene and palaeogene, respectively (see also japsen et al. 2012b). in south-east brazil, an extensive, low-relief erosion surface known as the japi surface, characterises much of the landscape at an elevation of 1200–1300 m a.s.l., but it locally reaches elevations of 2 km a.s.l. (e.g. almeida & carneiro 1998 (and references therein); riccomini et al. 2004). according to almeida & carneiro (1998) 16°s 15°s 14°s 13°s 40°w41°w42°w belo horizontebelo horizonte rio de janeirorio de janeiro salvadorsalvador elevation (m a.s.l.) hs hs hs hs hs ls ls ls hs hs 55 km55 km hs hs hs ls ls ls 2000 1800 1600 1400 1200 1000 800 600 400 200 fig. 65. altitude between chapada diamantina in the north and planalto da conquista in the south with interpretation of two peneplains, the higher surface (hs) and the lower surface (ls). the ls is between 300 and 400 m a.s.l. and dips slightly eastward. sharp, erosional, winding escarpment separates the ls from hs. the hs is particularly well preserved on planalto da conquista (c. 900 m a.s.l.) where the surface covers an area of c. 18,000 km2 within the map frame; note the wide and shallow valley that trends towards east. the hs is also clearly defined on chapada diamantina (c. 1200 m a.s.l.). location shown in fig. 64. modified from japsen et al. (2012b). 104104 the japi surface truncates a number of well-dated upper cretaceous intrusive complexes but not maastrichtian complexes, leading these authors to conclude that the formation of the japi surface was completed before the end of the late cretaceous. 6.4.3 studies focused on thermochronology gallagher et al. (1994, 1995) presented apatite fissiontrack analysis results from basement and sediment samples from the margin of south-east brazil. their fission-track ages broadly increase inland from between 60 and 90 ma on the coastal plain to >300 ma in the hinterland, and they concluded that the data reflected protracted denudation since the opening of the south atlantic. they based their analysis on what they considered to be the empirical scarp retreat model of gilchrist & summerfield (1990; see fig. 57), and concluded that their fission-track data were broadly consistent with that model. gallagher et al. (1998) noted that the youngest apatite fission-track ages along the south-east brazilian margin are significantly younger than the age of the rifting that they assumed led to the formation of the margin, and thus that substantial amounts of post-rift denudation had occurred. franco-magalhães et al. (2010) presented young apatite fission-track ages (66–6 ma) from the south-east brazilian margin that they interpreted as reflecting post-rift activation of the margin in late cretaceous, eocene and miocene times. however, they did not discuss how these very young ages relate to the truncated strata of ordovician to jurassic sediments and lower cretaceous flood basalts that are prominent in their study area. nor did they discuss their results relative to the post-rift japi surface that characterises the planaltos of their study area (cf. almeida & carneiro 1998). further, they did not discuss their apatite fission-track results in terms of timing and magnitude of discrete exhumation episodes and how this relates to formation of the japi surface relative to burial and exhumation of their study area. significant post-rift erosion has also been inferred from apatite fission-track studies of outcrop samples along the margin of north-east brazil even at considerable distance from the coast (harman et al. 1998; turner et al. 2008; morais neto et al. 2008). harman et al. (1998) identified two main phases of cooling, the first in fig. 66. landscapes and interpretation of peneplains in north-east brazil (bonow et al. 2009; japsen et al. 2012b). a: the higher surface (hs) in chapada diamantina, here at 1200 m a.s.l. with wide shallow valleys. b: the erosional chapada diamantina escarpment between the hs and lower surface (ls). the ls is here at c. 700 m a.s.l. c: the ls with wide and shallow valleys north-west of salvador. the surface is here at 250 m a.s.l. photo locations shown in fig. 64b. a b c 105 the early cretaceous (c. 130 ma) and the second in the late cretaceous (60 to 80 ma). harman et al. (1998) estimated a total cooling to the present-day along the continental margin of c. 100°c since 130 ma, and a cooling in the interior of their study area since 80 ma of 50–70°c. turner et al. (2008) also identified two major cooling episodes which began in the intervals from 110 to 100 ma (from palaeotemperatures higher than 75°c) and 40 to 10 ma (from palaeotemperatures between 40 and 80°c). morais-neto et al. (2008) interpreted their data in terms of two dominant episodes of cooling that began in the intervals from 100 to 90 ma (from palaeotemperatures higher than 80°c) and after 20 ma (from palaeotemperatures between 45 and 85°c), but they also reported evidence for an intermediate event in the interval from 65 to 50 ma. despite the differences in the assessments of the timing between these studies, it is apparent that rocks that are now at the surface in north-east brazil have cooled by 50°c or more since the mid to late cretaceous. combined with the presence of cretaceous sedimentary units in the region, these results imply that kilometrethick deposits once covered the rocks now exposed on present-day plateau surfaces, and that these deposits have been removed since the cretaceous. consequently the present plateau surfaces must be post-cretaceous erosional features. 6.4.4 integration of approaches the studies described in section 6.4.3 were focused primarily on interpretation of thermochronology data, without taking much account of other constraints. in contrast, cobbold et al. (2001), cogné et al. (2011, 2012) and japsen et al. (2012b) published studies of the postbreakup development of the brazilian margin based on integration of thermochronological data and geological observations, and in the case of japsen et al. (2012b) also with analysis of large-scale landforms. these studies lead to very different conclusions relative to the earlier thermochronological studies. cobbold et al. (2001) studied a variety of geophysical and geological observations as well as apatite fissiontrack data from the obliquely rifted margin of south-east brazil (20–27°s), and documented late cretaceous, eocene and neogene reactivation of older structures, attributed to the combined effects of far-field stresses and hot-spot activity. in particular, cobbold et al. (2001) found that the coastal mountains underwent block faulting and uplift in the neogene. cogné et al. (2011, 2012) studied the elevated margin of south-east brazil (22–24°s) where upper cretaceous/ palaeogene intrusions provide good evidence for postbreakup tectono-magmatic activity, and focused particularly on the cenozoic onshore basins in the area; e.g. the taubaté basin (cf. riccomini et al. 2004). they did not, however, include analysis of the landscape in their work, e.g. the extensive japi surface (almeida & carneiro 1998), and were thus unable to judge when the margin was close to sea level (compare figs 67 and 68). modelling of the thermochronological data together with geological observations led cogné et al. (2011, 2012) to conclude that three periods of post-rift accelerated cooling had affected the margin; during the late cretaceous, the palaeogene and the neogene. the areas near the cenozoic basins also experienced a period of burial prior to the neogene; see fig. 67. cogné et al. (2012) summarised that the postbreakup evolution of south-east brazil reflects a combination of structural inheritance, magmatic activity and plate-wide stress, leading to post-rift episodic uplift, rather than erosion of rift-related, uplifted relief. japsen et al. (2012b) reported the outcome of an integrated study of landscape development and thermo-tectonic evolution of the epcm in north-east brazil, focusing on the early cretaceous rtj rift and the extensive plateaus in the hinterland (fig. 64; mainly 10–15°s; cf. magnavita et al. 1994). bonow et al. (2009) and japsen et al. (2012b) analysed the landforms in the study area and identified two surfaces of low, relative relief and regional extent, referred to as the higher surface (hs) and the lower surface (ls) as discussed in section 6.4.2; cf. figs 65, 66. japsen et al. (2012b) reported afta data from outcrop samples from north-east brazil and from samples of sediment from boreholes down to 5.3 km below ground level in the rtj rift (fig. 64), and defined nine regional cooling episodes from a synthesis of thermal history solutions derived from afta data in all samples (table 4). the events date back to the palaeozoic, but japsen et al. (2012b) focused on the (post-rift) cretaceous and cenozoic development. the afta data define cooling episodes which began at c. 120 ma and in the intervals from 80 to 75, 48 to 45 and 18 to 15 ma (aptian, campanian, eocene and miocene cooling episodes). an albian event (beginning between 110 and 105 ma) is only recognised in samples east of the rtj rift. the afta and vr data from wells in the rtj rift show that the syn-rift sequences began to cool from their 106106 maximum post-depositional palaeotemperatures in the campanian, followed by two further cooling phases in the eocene and miocene (fig. 69; japsen et al. 2012b). analysis of the variation of palaeotemperature with depth in the wells shows that heat-flow conditions were close to those of the present day during all three episodes, implying that cooling in all three episodes was due to exhumation. campanian palaeotemperatures were attributed to additional (post-rift) burial by c. 3 km of post-rift section, while eocene and miocene palaeotemperatures indicate burial by c. 2 km and c. 1 km of additional section, respectively. the conclusions of earlier studies described above (harman et al. 1998; turner et al. 2008; morais neto d. late neogene to present day c. late palaeogene to early neogene serra da mantiqueira area tertiary basins area serra do mar coastal area b. palaeogene present surface level a. late cretaceous basement and possibly younger rocks 0 1 el ev at io n (k m a .s. l.) nw se palaeogene sediments of taubaté basin removed section of palaeogene–neogene sediments of taubaté basin inherited shear zone and faults apatite fission-track sample phase of exhumation fig. 67. the post-rift development of the south-east brazilian margin (23°s) based on thermochronological data and thermal history modeling according to cogné et al. (2012). a: in the late cretaceous, a considerable rock column covered the present-day surface prior to a first phase of post-rift uplift and erosion. b: during the early palaeogene, a second phase of cooling/exhumation occurred only within the area of cenozoic basins. c: during the late palaeogene and early neogene, the area around the cenozoic basins was buried under sediments. d: from the late neogene until the present day, a third phase of uplift and erosion affected the entire area leading to the formation of the present-day topography. cogné et al. (2012) concluded that the high topography along the margin of south-east brazil is the consequence of post-rift episodic uplift, rather than of erosion of rift-related, uplifted topography. sacek et al. (2012) presented a different conclusion involving no post-rift activation as discussed in the text. onset of cooling (ma) stratigraphic interval areas/rocks with cooling episode identified 450–410 ordovician–devonian * 320–300 carboniferous * 230–220 late triassic * 180–170 middle jurassic * c. 120 aptian most basement samples 110–105 albian only samples in restricted area 80–75 campanian sediment and basement samples 48–45 eocene deep well samples 18–15 miocene sediment and basement samples table 4. intervals for onset of cooling episodes in north-east brazil intervals defining the beginning of episodes of regional cooling based on afta data in 131 samples from outcrops and boreholes in north-east brazil (japsen et al. 2012b). * pre-cretaceous episodes recognised only in restricted areas where effects of more recent events (particularly in the aptian) are low enough to preserve evidence of earlier history. the early episodes likely affected much of the region. 107 et al. 2008) are broadly compatible with the results of japsen et al. (2012b). all studies emphasise the significance of regional cooling in the late cretaceous, even though earlier studies did not resolve the full complexity of the variation in the thermal history across the region. the greater detail obtained by japsen et al. (2012b) was possible because they also analysed samples from deep wells in the rift. the three main cooling episodes identified by japsen et al. (2012b) in north-east brazil match the late cretaceous, palaeogene and neogene events identified by cogné et al. (2012) in south-east brazil. japsen et al. (2012b) presented a synthesis of geological data, stratigraphic landscape analysis and palaeothermal and palaeoburial data, defining a four-stage history of post-rift episodes of burial, uplift and exhumation that shaped the atlantic margin of north-east brazil (fig. 68): (a) after early cretaceous break-up, the margin underwent burial beneath a thick sedimentary cover; (b) uplift episodes in the campanian and eocene led to almost complete removal of these deposits and formation of a large-scale, low/relief erosion surface (the higher surface, hs); (c) the hs was deeply weathered and finally reburied at the oligocene–miocene transition; and (d) miocene uplift and erosion produced a new, lowerlevel peneplain (the lower surface, ls) by incision of the uplifted and re-exposed hs. japsen et al. (2012b) noted that the uplift phases in brazil were synchronous with uplift phases in africa and the andes (fig. 70). this four-stage model agrees with the results of previous studies which found that the rtj rift was buried below a kilometre-thick cover prior to exhumation (magnavita et al. 1994), that the plateau surface (hs) in north-east brazil was fully developed by the end of the palaeogene (king 1967; valadão 1998), that the oligo– miocene transition was characterised by subsidence of the coastal regions throughout brazil (viana et al. 1971; schobbenhaus & brito neves 2003; rossetti et al. 2013), that plateaux elsewhere in north-east brazil are covered 2 1 0 0 hshs ls ls -1 -2 -3 -4 -5 d. present planalto do maracáschapada diamantina rift basin sea sea sea hs hs 0 -1 0 -1 -2 -3 c. c. 17 ma hs hs 0 0 -2 -1 b. c. 30 ma ? ? ? elevation (km a.s.l.) -1 -2 0 0 -1 -2 -3 -4 present hs level a. c. 78 ma w e oligocene – mid-miocene (16 ma) / pliocene (present) sediment mid-campanian – eocene sediment cenozoic laterite post-rift sediment, albian – mid-campanian rift sediment, berriasian–aptian pre-rift sediment, jurassic and older precambrian basement higher surface (hs) lower surface (ls) base of slide 75 km fig. 68. burial and exhumation history along a profile in north-east brazil across chapada diamantina and the early cretaceous rift basin, based on a topographic profile (b–b’ in fig. 64) and a geological cross section offshore (menezes & da silva milhomen 2008). a: c. 78 ma: campanian maximum burial of the lower cretaceous synrift sequence below a cretaceous cover that most likely extended over the basement terrains from the atlantic margin into the sanfranciscana basin (fig. 62). b: c. 30 ma: final formation of the higher surface (hs) by erosion to base level as a peneplain with deep weathering profiles and laterites after campanian and eocene phases of uplift and erosion. major sliding offshore took place after the campanian and eocene uplift events (cobbold et al. 2010). c: c. 17 ma: oligo–miocene burial of the interior highlands and of the coastal zone. d: present: following 1) miocene uplift which caused re-exposure of the hs and formation of the lower surface (ls) by river incision and valley widening with weathering, and 2) minor uplift in the quaternary which led to incision below the lower surface and to formation of a coastal plain. timing of events of cooling and exhumation from afta are listed in table 4. modified from japsen et al. (2012b) 108108 fig. 69. palaeothermal data from two wells defining burial and exhumation histories in the recôncavo-tucano-jatobá rift (location shown in fig. 64a). a: apatite fission-track analysis (afta) parameters for samples from wells a and b (drilled in the tucano and recôncavo basins, respectively) plotted against depth and present-day temperature. the black line in the left panel shows the increasing stratigraphic age with depth. coloured lines show the predicted patterns of fission-track age and mean track length for apatites containing <0.1, 0.5, 1.0, and 1.5 wt% cl from the default thermal history (dth). the default thermal history was derived from the preserved sedimentary section and the present-day thermal gradient calculated from corrected borehole temperatures (bht) in each well. the fission-track ages decrease systematically with depth. at depths greater than 2 km, the ages are much younger than the values predicted from the default thermal history. this shows that the sampled units have been hotter in the past (green et al. 2002). the pattern of decrease in fission-track age with depth is characteristic of a section that has undergone major cooling, with the ‘break-in-slope’ at a depth of c. 2.5 km representing the transition between partial and total annealing of fission tracks formed prior to the onset of cooling (fig. 21). the corresponding fission-track age of c. 75 ma represents the onset of exhumation. b: palaeotemperature constraints vs. depth in wells a and b for the campanian, eocene, and miocene palaeothermal episodes (table 4). drilled stratigraphy for each well is shown to the right. constraints for each episode in both wells define linear profiles, sub-parallel to the present-day temperature profile, characteristic of heating predominantly due to deeper burial. vr: vitrinite reflectance. c: ranges of amount of removed section and palaeogeothermal gradients (banana-shaped areas) required to explain palaeothermal profiles in wells a and b within 95% confidence limits. limits on the amount of removed section were also estimated from sonic data from sandstone units (rectangular areas). interpretations based on constant geothermal gradients corresponding to present-day conditions are also indicated, with palaeothermal and palaeoburial (based on sonic data) approaches giving highly consistent results in both wells (cf. japsen et al. 2007a, 2012b). the presentday temperature profile for well b is based on corrected bht data and temperatures revised on the basis of the afta data. modified from japsen et al. (2012b). mean length (µm) d ep th (k m ) fission-track age (ma) 0 1 2 3 4 5 0 100 200 300 400 30 40 50 60 70 80 90 30 40 50 60 70 80 90 stratigraphic age 4 8 12 16 cl cl cl fission-track age with error bar unconformity well a well b prediction from dth increasing chlorine content well a eocene profile eocene profile campanian profile campanian profile miocene profile miocene profile4 1 2 0 3 d ep th (k m b el ow g ro un d le ve l) 4 1 2 0 3 d ep th (k m b el ow g ro un d le ve l) 0 40 80 120 160 200 temperature (°c) well b 0 40 80 120 160 200 temperature (°c) corrected bht afta sample horizon maximum palaeotemperature from vr core measurement outlier range of palaeotemperatures from afta present-day thermal gradient of 15°c/km post-rift (115–110) rift (144–123) pre-rift (147–146) sediment (ma) geothermal gradient (°c/km) maximum likelihood solution parameter ranges from afta 95% confidence limits sandstone sonic data sandstone sonic data re m ov ed se ct io n (k m ) re m ov ed se ct io n (k m )well a well b campanian campanian eocene eocene miocene miocene constant heat flow: 15°c/km 5 5 4 3 2 1 0 10 15 20 25 c. 3.0 km c. 2.5 km c. 2.0 km a b c 5 4 3 2 1 0 te m pe ra tu re (° c ) present-day thermal gradient of 20°c/km c. 3.5 km c. 2.5 km c. 1.5 km geothermal gradient (°c/km) 10 15 20 25 30 constant heat flow: 20°c/km 69 109 with non-fossiliferous, continental sediment (e.g. the palaeogene serra do martins formation on borborema plateau; morais neto et al. 2008) and that uplift affected much of brazil in the miocene (king 1967; valadão 1998; cobbold et al. 2001). aspects that have remained unrecognised in previous studies, however, are the absolute timing of the episodes of uplift and the magnitudes of the burial and exhumation episodes that followed early cretaceous rifting and eocene–oligocene peneplanation. 6.4.5 continuing controversy while the various thermochronology studies discussed above led to differing conclusions in regard to the evolution of the continental margin of brazil, they all consistently provide evidence of kilometre-scale denudation across a wide region. where synor post-rift sedimentary remnants are preserved, this implies deposition of significant thicknesses of sedimentary cover that has been subsequently eroded. but some authors regard such interpretations as unrealistic. peulvast et al. (2008) discussed the apparent conflict in north-east brazil, between the amount of cover that has been removed following atlantic break-up based on landscape studies and amounts derived from apatite fission-track data. the idea that a substantial thickness of section could have been deposited and then eroded was regarded as unrealistic, despite evidence from apatite fission-track data (harman et al. 1998; morais neto et al. 2006, 2008) showing that samples now at surface were at 70–85°c in the late cretaceous. peulvast et al. (2008) found that the thickness of the eroded post-rift cover indicated by these studies by far exceeded their estimates based on the exposed strata, and preferred to dismiss the apatite fission-track results as either reflecting lowtemperature artefacts in fission-track annealing models or effects of cenozoic magmatism. but the results presented by japsen et al. (2012b) document that the high late cretaceous palaeotemperatures revealed by apatite fission-track data from across north-east brazil are not artefacts of the annealing models since they are confirmed by vr data and also by interpretation of sonic data from deep boreholes in the rtj rift (fig. 69c). an explanation of post-rift palaeotemperatures in and around the rtj rift in terms of cenozoic magmatism can be ruled out, since none has been reported from that area. in addition, the palaeothermal data from the deep boreholes in the rtj rift show that the palaeogeothermal gradient has been low (10–30°c/km) since the campanian (fig. 69c). there are also diverging views between studies of the burial and exhumation history of south-east brazil. as referred to above, cogné et al. (2012) combined apatite fission-track data with geological observations to conclude that the high topography along the margin of south-east brazil is the consequence of post-rift episodic uplift, rather than of erosion of rift-related uplifted topography. furthermore, these authors showed that the basement rocks now exposed at the surface were buried below a kilometre-thick cover in the late cretaceous, and that they remained buried until the late neogene outside the taubaté basin. in contrast, sacek et al. (2012) used numerical modelling to investigate the same segment of the brazilian margin as that studied by cogné et al. (2012) but came to conflicting conclusions, as they found that the present-day morphology along that margin can be explained as resulting from rift-related vertical motions alone, without requiring significant post-rift “rejuvenation” (their quotation marks). sacek et al. (2012) interpreted their modelling results to show that the combination of rapid erosion of the coastal (rift) escarpment and differential subsidence of the margin favours the emergence of a secondary bulge and its establishment as the main, ‘permanent’ drainage-divide escarpment, while the coastal escarpment progressively disappears. sacek et al. (2012) did not consider the possibility that rocks now exposed away from an initial escarpment have been more deeply buried at any time since rifting and break-up. reconciliation of thermochronological data with the geological record and landscape evidence – particularly in terms of recognising the former presence of covers that are now removed – remains a matter of controversy in discussion of the development of the brazilian margin. as a result, the extent, origin, age and importance of the brazilian planaltos remain controversial in the tectonic analysis of the margin. from the above review, it seems clear that the brazilian plateaux and their present elevation are post-rift features, and that they were graded towards the general base level, which in the case of the post-rift development of northeast brazil was determined by sea level of the adjacent atlantic ocean (e.g. king 1967; almeida & carneiro 1998; valadão 1998; zalán & oliveira 2005). yet many studies of the development of the brazilian margin do not integrate the presence of these large-scale peneplains into their analysis (e.g. gallagher et al. 1998; franco110110 magalhães et al. 2010; cogné et al. 2012; sacek et al. 2012). in contrast, almeida & carneiro (1998) stressed that the dating of the extensive japi surface as a stratigraphic marker is of crucial importance for understanding the development of the brazilian margin. this statement emphasises the importance of the accurate identification of such erosion surfaces. for example, can the japi surface at c. 1200 m be confidently correlated with the flat summits at c. 2 km a.s.l. as suggested by almeida & carneiro (1998), and is it the same surface that truncates pre-maastrichtian but not maastrichtian intrusive complexes? to our knowledge, no such mapping in brazil has yet been made. careful mapping of the extent of erosion surfaces such as the japi surface becomes a cornerstone for evaluating the relation between the surface and geological markers and hence for incorporating the presence of elevated peneplains in our understanding of the development of epcms. 6.4.6 summary in contrast to studies which assumed that the atlantic margin of brazil was uplifted at the time of rifting and remained as a positive region since, integration of thermochronology with observations of geology and landforms have shown that the margin subsided after breakup and was buried by up to 3 km of post-rift sediment sediments ng pg lc ec j ng pg lc ec j ls hs lps ups uncertain age hiatus continental reburial reburial marine post-rift br ea kup rift pre-rift unconformity br ea kup br ea kupri fti ng ri fti ng ri fti ng pe ru via n in ca ic q ue ch ua n d rif t s to p a. nuussuaq basin b. rtj rift c. rio muni basin maximum burial post-rift thermal subsidence formation of peneplain ng pg lc ec j age (ma) 0 10 20 30 40 50 60 70 80 90 100 110 120 130 140 150 sr d ro p sr d ro p onset of exhumation from afta/ geomorphology tectonic events drop in spreading rate fig. 70. comparison of timing of post-rift uplift events, formation of peneplains, and tectonic episodes along margins in the atlantic domain. a: nuussuaq basin (c. 70°n), west greenland (japsen et al. 2006). the upper planation surface (ups) defines the plateau whereas the lower planation surface (lps) is a system of palaeovalleys (figs 34, 35; bonow et al. 2006a, b). events of subsidence and uplift along the west greenland margin, prior to break-up are shown in fig. 38. b: recôncavo-tucano-jatobá rift (rtj; c. 12°s), north-east brazil. the higher surface (hs) defines the plateau. the lower surface (ls) is an extensive low-relief surface that has developed at the expense of the hs (figs 65, 66). stratigraphy is after viana et al. (1971) and silva et al. (2007). main phases of andean orogeny: peruvian, incaic, and quechuan (pardo-casas & molnar 1987; cobbold et al. 2001, 2007). c: rio muni basin (c. 2°n), west africa (turner et al. 2008); conjugate margin to north-eastbrazil. the colours of the horizontal bands indicate onset of uplift events and the interpreted correlation between events on the conjugate margins in north-east brazil and west africa. the present high topography with plateaux in north-east brazil and west greenland were formed millions of years after break-up when regional peneplains were uplifted and dissected during the neogene. changes in plate motion after chalmers & pulvertaft (2001) and torsvik et al. (2009). afta: apatite fission-track analysis. ec: early cretaceous. j: jurassic. lc: late cretaceous. ng: neogene. pg: palaeogene. sr: spreading rate. modified from japsen et al. (2012b). 111 (magnavita et al. 1994; morais neto et al. 2006, 2008; japsen et al. 2012b). this cover was subsequently almost entirely removed in a series of episodes of uplift and erosion that first led to formation of an extensive erosional surface near sea level and then to its uplift and dissection, leaving the present-day plateau remnants. these studies thus document that the present-day high topography along the brazilian epcm formed long after break-up. this supports conclusions from early studies based purely on observations in the field (e.g. king 1956a; almeida & carneiro 1998). a significant difference is that the early studies did not consider the possibility of burial below post-rift sedimentary cover which was subsequently removed. the evolution of the north-east brazil epcm defined above is similar in many respects to that defined for west greenland in chapter 5, and as also inferred for southern africa (section 6.2) and south-east australia (section 6.3). we therefore suggest that this style of development, involving post-breakup subsidence and subsequent uplift and erosion in a series of episodes, describes the evolution of many epcms and may indeed provide a general description of their evolution. 6.5 other areas comments similar to those in relation to southern africa, south-east australia and brazil, discussed above, also apply to a number of other areas where apatite fission track and geomorphology have traditionally proved difficult to integrate. nowhere is the inconsistency between the two approaches more pronounced than in northern australia, where landforms regarded by some as having remained unaffected from early palaeozoic times or even earlier have yielded much younger apatite fission-track ages, suggesting kilometre-scale denudation over the last 300 million years (belton et al. 2004). numerous geomorphological studies (as reviewed by belton et al. 2004) have contributed to the idea that australia preserves some of the world’s oldest landscapes, reflecting the tectonic stability of the cratonic cores of the continents. however, extensive mapping of apatite fission-track data across the continent (gleadow et al. 2002) has shown an absence of apatite fission-track ages in excess of 500 ma, suggesting that the continent has experienced a much more tectonically active history than previously supposed. in the yilgarn block of western australia, weber et al. (2005) reported fission-track ages generally between 200 and 300 ma in samples from present-day outcrops, refuting the idea of the long term stability of this archean craton. indeed, the results presented by weber et al. (2005) actually show that the supposedly stable shield was reexposed from below a sedimentary cover and reburied during the permian, as discussed in chapter 4. the tendency to see topography as old is well illustrated by persano & dobson (2009) who sought to provide an answer to bob dylan’s musical query about erosional timescales. they presented an interpretation of apatite fission-track and apatite (u-th)/he data from two vertical profiles in north-west scotland in terms of long-term, post-caledonian, monotonic cooling reflecting denudation over 400 ma. however, the presence of late devonian volcanics and sedimentary units near one of their transects and triassic sandstones at outcrop over the wider region shows that these areas were close to the surface at various times in the post-caledonian history. holford et al. (2010) showed that taking these sedimentary remnants into account results in a much more active history involving multiple episodes of burial/deposition and uplift/exhumation than that suggested by persano & dobson (2009). northern england provides an example of a situation where apparently incompatible results from afta and landscape/geological investigations have been brought into a consistent regional framework. in initial interpretations of afta data from the area (green 1986, lewis et al. 1992), samples which cooled from greater than 110°c in the paleocene (c. 60 ma) were interpreted in terms of burial by >3 km of section that was removed during cenozoic denudation. such thicknesses of former cover were regarded by many as unrealistically high (e.g. holliday 1993). however, later, more detailed studies involving afta data in samples over vertical sections from both wells and outcrop sections revealed that palaeogeothermal gradients were seriously underestimated in previous studies, and use of more appropriate values around 50 to 60°c/km as defined from measured afta and vr data suggested much lower amounts of eroded section (up to c. 1.5 km), allowing the two approaches to be readily reconciled (green 2002). while the focus of discussion in this chapter has been to highlight inconsistencies in the interpretation of landscape and thermochronology studies, it should also be noted that some examples exist where the two approaches have shown a much greater degree of consistency, including studies of south-east asia (schoenbohm et al. 2004) and corsica (kuhlemann et al. 2005). the key to successful integration of the different approaches in 112112 these areas lies in data from both approaches being given equal weight, with the evidence inferred from landscape studies used to provide prime constraints on the evolution, similar to the studies in west greenland and brazil discussed here. we suggest that an approach similar to that outlined in chapter 5, involving detailed thermochronology analyses in carefully collected samples, with data interpreted within a framework constrained by geological observations and stratigraphic landscape analysis, will ultimately lead to reconciliation of results from different approaches in southern africa, south-east australia and elsewhere, and to an improved understanding of the development of epcms. similar comments apply to cratonic regions. as shown in chapter 3, re-exposed peneplains with characteristic relief (the flat sub-cambrian and the hilly sub-mesozoic peneplains) in basement rocks are of major importance in the present landscape of south sweden and this despite several quaternary glaciations. such re-exposed forms also make up much of the relief farther north (lidmarbergström 1995, 1996; lidmar-bergström et al. 2013). similar observations are also made on the canadian shield with a flat sub-ordovician surface (ambrose 1964) and a hilly sub-mesozoic surface (lidmar-bergström & jansson 2005). however, such observations are seldom acknowledged in recent studies dealing with long-term erosion. integration of stratigraphic landscape analysis and thermochronology in regions such as these will provide an improved understanding of subsidence and uplift of shield areas, previously widely regarded as stable (cf. the north australian and western australian examples reviewed above). 6.6 numerical modelling of epcm development in recent years, considerable effort has been devoted to development of numerical models describing the geomorphological development of rifted margins (e.g. van der beek et al. 1994, 1995; braun & van der beek 2004; sacek et al. 2012). these models commonly seek to explain the form of epcms as a result of erosion acting on already elevated rift flanks formed at high elevations that were already extant at the time of break-up, within a framework of progressive denudation. they further assume that denudation is focused seaward of the coastal escarpment, with high-level plains being much less affected, in similar fashion to the conceptual models illustrated in fig. 57. in this way, the need for a tectonic component of uplift associated with (or following) rifting and break-up is eliminated (although the high elevation remains unexplained), and the morphology is then explained purely in terms of erosional processes (moderated by changes in climate) and the corresponding isostatic response, within a framework of progressive emergence and monotonic cooling. an extension of these studies is the three-dimensional, thermal-kinetic modelling package pecube, the current status of which, together with a large number of applications, is reviewed by braun et al. (2012). this allows prediction of vertical and horizontal patterns of ages from various thermochronological methods, but is focussed towards upward crustal movements involving progressive denudation of basement regions. as discussed above, evidence from southern africa, south-east australia, brazil and other regions, in addition to west greenland (chapter 5), defines a different style of behaviour, revealing major post-breakup burial and subsequent exhumation, extending hundreds of kilometres inland from the escarpment, often in multiple episodes. so the relevance of the numerical models described above to real world situations is far from clear. such models include no tectonic input to the development of the margin. however, in regions such as south-east australia, southern africa, brazil and west greenland, where additional burial is required to produce the palaeotemperatures prior to the onset of cooling defined from thermochronological data, it is also necessary to explain the transition from subsidence/burial to uplift/exhumation. in such circumstances, a tectonic input would appear to be mandatory. in summary, models of epcm development based on continuous denudation do not provide an accurate description of the processes involved. new approaches are required which include post-breakup subsidence and burial, in addition to subsequent denudation and uplift, in order to provide an accurate description. 6.7 concluding remarks in those situations discussed above where conclusions derived from thermochronology conflict with those from landscape analysis, often little attempt has been made to integrate both approaches with equal weighting. more typically, proponents from each side have looked at the region with a mindset based on their own experiences 113 and ideas, and failed to take into account what appears to be conflicting evidence from another point of view. common examples include the adoption of constant heat flow in thermochronological studies despite abundant evidence that this is untenable, failure to acknowledge the information provided for example by erosion surfaces (peneplains) and valley generations in landscape studies and their relevance to the interpretation of thermochronology data, and attempts to ‘shoe-horn’ data into compliance with simple conceptual models. the omission of any discussion of apatite fission-track data in many landscape studies, as discussed earlier, amounts to effectively discounting the relevance of the apatite fission-track approach. we suggest that if observations are based on reliable science, then a satisfactory reconciliation must be possible, and that all that is required to achieve this is an attitude that is receptive to novel outcomes, rather than demanding that results conform to expectation and accepted models. the next chapter discusses various mechanisms which might explain the episodic development of epcms described in the preceding chapters. 114114 7. what processes drive the formation of epcms? plate-tectonic theory accounts well for mountain ranges on continents formed at convergent or collisional plate margins, whether they are due to subduction of an oceanic plate below a continental plate or to the collision of two continents. plate-tectonic theory does not, however, account well for the mountain ranges discussed in this paper; those reaching 1–2 km a.s.l. found inland from many passive (extensional) continental margins (e.g. japsen & chalmers 2000; japsen et al. 2012a). such epcms have been recognized for many years (chapter 1), but many explanations in the literature for their presence are unconvincing. the special properties of only one epcm are commonly used to try to account for its existence, despite their common aspects (e.g. asymmetrical topography and high plains), while others make assumptions about properties of epcms that are not in accordance with observations. we review here some of the more popular hypotheses. 7.1 hypotheses that apply to one or only a few epcms some hypotheses concerning the formation of individual epcms rely on some particular property that is not shared by others. for example, much of the norwegian epcm coincides with the scandinavian part of the caledonian orogenic belt. mountain ranges formed by continental collision develop deep roots that hold their topography uplifted by isostasy (e.g. watts 2001, pp. 339–361). nielsen et al. (2009) assumed that the norwegian mountains are merely the eroded remnants of former caledonian topography and that they are held uplifted by a remnant crustal root. this hypothesis has now been tested and disproved (stratford & thybo 2011; ebbing et al. 2012; maupin et al. 2013). no root is present beneath the northern scandes (ebbing et al. 2012) and, while there appears to be a small root under the southern scandes (stratford & thybo 2011), it is not sufficient to support the main mountain range isostatically and is displaced from it. there is also good evidence that the scandinavian caledonides collapsed in the devonian (e.g. dewey et al. 1993) while the branch of the caledonian orogeny under northern poland and germany is buried beneath palaeozoic and younger sediments. see lidmar-bergström & bonow (2009) and chalmers et al. (2010) for a more comprehensive discussion of nielsen et al. (2009). similarly, pedersen et al. (2012) suggested that the high topography in east greenland is an erosional remnant of topography originally developed during the caledonian orogeny. pedersen et al. (2012) used apatite fission-track data in an attempt to demonstrate a monotonic cooling history for that region, representing continual uplift and denudation, since the end of the caledonian orogeny. however, as discussed by japsen et al. (2013a) the geological record of east greenland shows that the caledonian mountains there had collapsed by the late palaeozoic (see section 8.2) and the region was, at least partially, buried below deltaic and marine carboniferous sediments. surlyk et al. (1984) showed that a peneplain had formed at sea level over much of northern east greenland by the mid-permian, and the area was again reburied under mesozoic sediments. haller (1971, p. 321) summarised the demise of the caledonian fold belt in the following way: “in central east greenland, which seems to have been the heart of the ancient mountain belt, the vigorous late movements lasted into the permian. the molasseladen caledonian mountain stumps then became completely bevelled and during the late permian the sea advanced from the east. this transgression marked the end of the caledonian orogeny. hence all subsequent events are to be included in the post-caledonian history”. a similar style of long-term evolution was suggested for the highlands of south-east australia by lambeck & stephenson (1986) who suggested the present-day mountains were a residual of the palaeozoic lachlan fold belt, with the topography maintained since c. 250 ma by the isostatic response to erosional unloading. this model is inconsistent with thermochronological evidence of major early cretaceous denudation across the region as reviewed in section 6.2, and also fails to explain the presence of eocene, low-level, fluvio-lacustrine muds at c. 1.8 kilometres above sea level, while similar muds also occur in adjacent basins at depths of up to 2 km below sea level (holdgate et al. 2008), as discussed in section 6.3.4. another hypothesis that applies to only a few epcms is based on the observation that some of them coincide with, or are at least near, large igneous provinces (lips). it appears to be the case that, while mantle upwelling (plume emplacement) immediately prior to emplacement 115 of a lip causes uplift (e.g. dam et al. 1998; saunders et al. 2007), there is no evidence that such uplift persists after cessation of volcanism, as has been suggested by some authors (e.g. mackay et al. 2005). subsidence during and after volcanism greater than the immediately preceding uplift has, however, been demonstrated in some lips such as west greenland (section 5; japsen et al. 2005) and east greenland (brooks 2011; bonow et al. 2014, in press). in addition, while some epcms contain lips, others do not (e.g. south-east australia and norway), and the west greenland and baffin island epcms run continuously from a lip to a non-volcanic margin (chalmers, 1997; 2012). 7.2 permanently uplifted rift margins? as discussed in chapter 2, much of the literature (e.g. gallagher et al. 1998; brown et al. 2002; persano et al. 2006; swift et al. 2008; sacek et al. 2012) assumes that epcms are erosional modifications of rift margins that were uplifted at the time of rifting/onset of sea-floor spreading and have remained uplifted since. some attempts to account for epcms as geodynamic features have also assumed that they have been present since the time of rifting or break-up (e.g. weissel & karner 1989; gilchrist & summerfield 1990; chéry et al. 1992; ten brink & stern 1992; watts 2001). however, as weissel & karner (1989) showed and japsen et al. (2012a) discussed, rift margins will remain uplifted indefinitely only if there is little or no extension of the mantle below the rift margin at the time of rifting (fig. 71); i.e. when the extension approximates simple shear. understanding of rifts and rift margins has progressed enormously in recent decades. early debates on rift formation concentrated on two end-member models; pure shear (e.g. mckenzie 1978) and simple shear (e.g. wernicke 1985). understanding of tectonic controls on sedimentation within a rift has also developed from the early days of sequence stratigraphy (e.g. payton 1977), when it was assumed that the relative changes in sea level, necessary to generate the third and higher-order unconformities that form sequence boundaries, are eustatic (e.g. haq et al. 1987 and references therein). studies since then have shown that simple shear and pure shear are probably end members of a continuum of extensional styles, and that extensional style and amount can vary not only between the crust and mantle (cloetingh & ziegler 2009), but at different levels within the crust (e.g. depth-dependent stretching; davis & kusznir 2004). while it is undoubtedly the case that a number of factors, such as variations in the total net amount of oceanic crust and variations in the terrestrial ice budget, can cause eustatic sea level to vary, it has also become apparent that modest stress fluctuations on a regional scale can lead to relative sea-level changes of the same order (100–200 m) and on time-scale characteristics (c. 3 million years) for third order cycles (cloetingh et al. 1985). evidence has also been accumulating steadily that, in addition to the modest relative changes in sea level recorded by third and higher-order cycles, rifted basins and margins have been subjected to breaks in subsidence and sedimentation that last 10 million years and more (e.g. a. pure shear b. simple shear 100 km β=1 rift margin 0 60 d ep th (k m ) 120 β=1 rift margin β=2 crust and mantle β=2 mantle β=2 crust 100 km 0 60 d ep th (k m ) 120 continental crust lithospheric mantle oceanic crust fig. 71. cross-sections showing the difference between a: pure shear extension and b: simple shear extension. in a the amount of extension of the upper mantle and crust is the same at the same place. in b, extension of the upper mantle takes place much farther basinward than extension of the crust, shown by β = 2 being at different locations in the crust and mantle. post-rift cooling causes subsidence of the basin margin in a but not in b, so any flexural uplift of the margin produced during rifting will also subside in a but not in b. extensive modelling of subsidence of basins with low beta-factors (<2) by the hydrocarbon industry has shown that mckenzie’s (1978) pure shear model of rifting is a good first approximation for the development of these basins, although extension at high beta factors appears to take place as in b. 116116 ziegler & cloetingh 2004; cloetingh & ziegler 2009). the unconformities that result from these events are second order and unconformities of this type in the southern west greenland basins have already been described in section 5.6. extensive modelling of subsidence of marginal basins by the hydrocarbon industry has shown that mckenzie’s (1978) pure shear model of post-rift subsidence approximates the actual depositional record quite well in basins where extension factors are less than about 2, except where the sedimentary record is interrupted by second order unconformities. the subsidence curves predicted by the mckenzie (1978) pure shear model are offset across these unconformities (see e.g. fig. 10 in cloetingh & ziegler 2009). the offsets are not what would be expected if sedimentation merely ceased, but are more consistent with the erosion of a substantial thickness of sediment prior to renewed subsidence. if a rift forms by processes approximating simple shear, the lithospheric mantle underlying the proximal part of the rift (at extension factors less than about 2) extends less than the overlying crust (fig. 71b). the finite strength of continental crust causes the area just outside the rift to rise because of its flexural response (see e.g. watts 2001, pp. 286–339). the proximal part of the rift basin will not subside thermally after rifting ceases because the underlying lithospheric mantle is not warmer than it was pre-rifting, so the uplifted rift margin will not subside either, but will only be lowered by erosion. on the other hand, the lithospheric mantle underlying a rifted margin that extends by processes approximating pure shear extends by the same amount as the overlying crust. in this case the rift will subside because the thinned and therefore heated underlying lithospheric mantle cools and subsides, causing the overlying proximal rift to subside too (mckenzie 1978). as in the case of simple shear, the finite strength of continental crust causes the margin just outside the rift to rise because of its flexural response, but it will subside again as the rift subsides because of the flexural response to both the finite strength of the thermally subsiding rift and to loading by sediments deposited in the basin. these effects, together with erosion of the margin while uplifted, will cause the margin outside the rift to subside below sea level and sediments will transgress across the rift margin forming a so-called ‘steer’s-head’ geometry in cross-section (watts et al. 1982; white & mckenzie 1988; braun & beaumont 1989; roberts & yielding 1991). the width of uplifted area caused by flexural uplift of a rift margin is also generally much smaller than the observed widths of epcms. figure 72 shows the calculated widths of uplifts caused by flexural uplift using elastic thicknesses of 15 km and 30 km compared with cross-sections through actual epcms. it is readily apparent that the actual epcms are all much wider than the calculated flexural uplifts. much wider flexural uplifts can be generated by assuming larger elastic thicknesses, as great as 115 km (e.g. chéry et al. 1992; ten brink & stern 1992), but those are much greater than the elastic thickness calculated independently for most epcms (watts 2001, fig. 8.30). elastic thicknesses correspondred sea el ev at io n (k m a .s. l.) el ev at io n (k m a .s. l.) el ev at io n (k m a .s. l.) el ev at io n (k m a .s. l.) el ev at io n (k m a .s. l.) 2 0 2 0 2 0 2 0 direction of the ocean norway australia brazil 200 km w e we e w sw africa we 2 0 african side sw swne nearabian side 2 0 0 -2 2 a b te = 15 km te = 30 km 0 -2 2 200 km fig. 72. comparison of cross-sections of elevated, passive continental margins (epcms) with the expected amount of flexural uplift of a rift margin. a: cross-sections of five different epcms (all at the same horizontal and vertical scale). the sections are oriented so that the nearest ocean is to the left and the continental hinterland to the right. b: flexed rift margins with uniform stretching in the crust and lithospheric mantle and with elastic thicknesses of 15 and 30 km immediately after rifting, shown at the same scale as the cross-sections in a. the amount of uplift is somewhat smaller and the width of the uplifted area is substantially narrower than the real examples (that all represent post-rift situations). as the margins become re-attached to the rift once extension (rifting) ceases, cooling of the extended lithospheric mantle under the adjacent rift will cause the margins to subside along with the rift. modified after japsen et al. (2012a). 117 ing to that of the lithosphere could arise if the crust and lithospheric mantle were locked and extended together, i.e. in pure shear, but then the rift margin should subside and form a ‘steer’s head’ after cessation of rifting. it is difficult to envisage where such large elastic thicknesses might arise in the case of simple shear, because that failure mode requires a weak layer somewhere in the deep crust or upper mantle to allow extension of the crust and mantle to be decoupled. in that event, the elastic response of the crust and mantle should be independent and observed elastic thicknesses should be of the same order as the strong layer of the crust. extension to form rifts at extension factors of less than about 2 appears to be approximated by simple shear, interrupted by episodes of uplift that form second order unconformities. post-rift subsidence is close to the mckenzie (1978) model, and the rift margin should subside by flexural response to sediment loading to form ‘steer’s head’ geometry. since this appears to be how the proximal parts of most passive margins extend, ‘steer’s head’ geometry should be normal on a passive margin, and there should be no permanently-uplifted rift margin. 7.3 uplift by underlying hot and/or upwelling mantle a number of authors have proposed that epcms may be uplifted by mechanical support from an underlying mantle upwelling or plume or by isostatic uplift from lowdensity asthenosphere originating from such plumes. jones et al. (2012), for example, used the admittance, which is the ratio of the coherent parts of the gravity and topography (mckenzie & bowin 1976; watts 2001), to suggest that the uplifted margins around southern africa are supported by several mantle diapirs. only the longer wavelength part of admittance reveals mantle isostatic anomalies. shorter wavelength anomalies show isostatic anomalies arising from the crust and have been used to calculate elastic thicknesses of the lithosphere (watts 2001), but mckenzie & fairhead (1997) have shown that estimates of elastic thickness using calculations of admittance are erroneous in areas of smooth topography, because there is insufficient signal from the topography. much of the interior of southern africa is very flat, although at an altitude of between 1 and 2 km, whereas the topography of the margins is much more rugged. it might be that jones et al.’s (2012) calculations show edge effects of a much larger mantle diapir under the whole of southern africa which is lifting an area around 2000 km in diameter (ni et al. 2002). in any case, it is difficult to see why mantle plumes should preferentially arise under passive continental margins as suggested by jones et al. (2012). the margins on opposite sides of the south atlantic are both uplifted, but are moving apart. there is evidence for uplift of southern africa by a mantle diapir (ni et al. 2002), but there is no large-scale, free-air gravity anomaly across the south atlantic that suggests that any mantle diapir extends there. so any plume under eastern brazil would have to be independent of the african one, but would yet have to keep pace with the moving south american plate. this is special pleading. it might be that the presence of a passive margin could induce mantle convection. this idea has been investigated by king & anderson (1998) who showed that any induced convection in such a place in the presence of horizontal asthenospheric flow would cause downwelling of the mantle on the continental side of the passive margin, not upwelling. recently, rickers et al. (2013) published a high-resolution, s-wave velocity interpretation of the north atlantic region that showed a low-velocity layer beneath much of the oceanic lithosphere, consistent with the long-wavelength, bathymetric high of the north atlantic. the low-velocity layer extends locally beneath the continental lithosphere of the southern scandinavian mountains, the danish basin, part of the british isles and eastern greenland. rickers et al. (2013) argued that the spatial correlation between the low-velocity layer and uplifted regions suggested dynamic support by low-density asthenosphere originating from the iceland and jan mayen hotspots. rohrman & van der beek (1996) had earlier proposed that the norwegian mountains could be supported by a diapiric intrusion into the lithosphere of some anomalously warm asthenospheric material flowing from the icelandic hotspot. this hypothesis was tested by pascal & olesen (2008), who used heat-flow measurements over norway to show that there is no abnormally hot mantle under the norwegian mountains, and this observation would seem to suggest that rickers et al.’s (2013) modelling may not account for all of the factors uplifting the scandinavian mountains. estimates of the amounts of dynamically supported uplift of a continental margin are based on the amounts calculated from adjacent, abnormally shallow oceanic crust. extension of the calculations to the adjacent continents suggests that there should be up to a kilometre of dynamic support in onshore central east greenland, smaller amounts in onshore norway and the british isles 118118 and negligible amounts in west greenland and baffin island (e.g. steinberger 2007), the exact amounts depending on assumptions about the density and rheology of the continental crust and upper mantle. a kilometre of dynamic support is insufficient on its own to explain the heights of the east greenland epcm, and there is insufficient uplift even when the effect of dynamic uplift is combined with the calculated amount of flexed uplift caused by erosion of the fjords and glacial valleys (medvedev et al. 2013). japsen et al. (2013a), however, drew attention to the much higher elevation of east greenland compared to west greenland and suggested that this difference could be due to dynamic support in the east from the iceland plume. other epcms that probably contain elements of dynamic support from present-day mantle upwelling are the margins around the red sea (al-hajri et al. 2009). there is, however, no a priori reason to invoke presentday mantle upwelling under eastern australia, western india and eastern brazil if other explanations for the presence of epcms there can be found. 7.4 uplift due to compression passive continental margins are, by definition, produced under conditions of extensional stress. continental crust extends, thins and eventually breaks, and oceanic crust forms in the break and widens such that the now riven continental fragments move apart. the extended margin then cools and subsides and sediment is deposited onto the subsiding continental and oceanic crust. the lithosphere must continue to extend at the mid-ocean ridge, implying net extensional stresses there, but there is increasing evidence that the continental margins on either side of the ocean, and even oceanic crust, may come under compression after breakup (e.g. mcadoo and sandwell 1985; cloetingh & wortel 1986; papers in doré et al. 2002; cobbold et al. 2007; holford et al. 2011). the world stress map project (heidbach et al. 2008) shows that, where data exist, all epcms are under compression. a number of authors (e.g. løseth & henriksen 2005; pedoja et al. 2011; japsen et al. 2012a) have speculated that the uplift of epcms may be caused by compression, but have not proceeded to a quantitative model of how such compression may affect a passive continental margin. cloetingh & burov (2010 and references therein) have argued that continental crust in many parts of the world is under sufficient compression that sinusoidal, crustalscale folds have formed, and cloetingh et al. (2008) extended the discussion to epcms and suggested that they formed by the effects of compression. the folding described by cloetingh & burov (2010) cannot arise from strains originating in ridge push, because even the weakest continental crust is too strong. strain from ridge push amounts to around 40 mpa, though this can rise to around 100 mpa if a large mantle plume head underlies the ridge (bott 1993). compression stress sufficient to induce folding of continental crust depends on both the chemistry of the crust and its temperature regime. there is general agreement that the uppermost 12-20 km of continental crust consists predominantly of wet, quartz-rich lithologies, such as granite and granitic gneiss. weak lower crust can arise if the lower crust is also granitic or if it consists of more basic materials and the temperature gradient is high; so high that uppermost mantle would be melting and basaltic volcanism would be expected in such areas. continental crust could fold under an imposed stress of around 300 mpa if its strength resides primarily in the uppermost c. 15 km, but the resulting folds would have wavelengths of 100 km or less; much shorter than the typical 200–250 km wavelength reported by cloetingh & burov (2010). there would be no detachment between continental crust whose lower part consists of basic or ultrabasic rock and the upper mantle. in such a case, the entire lithosphere would have to fold as a unit. this would require compressional stresses approaching 2 gpa and would give rise to folds with a wavelength around 400–500 km. mcadoo & sandwell (1985) have reported folded oceanic lithosphere in the indian ocean that would require such stresses, where cloetingh & wortel (1986) model a concentration of compressional stress. it is, however, possible to fold continental lithosphere at much lower stresses if the lower crust is of intermediate composition. folds with wavelengths of the order of 200–250 km, as reported by cloetingh & burov (2010), can arise if the lower crust is of intermediate composition and is detached from the underlying mantle by a weak zone at the base of the crust (see e.g. burov 2009). to fold such crust requires stress of the order of 1gpa; stress of the same order as that required to drive an orogeny. it is therefore possible that stress originating from an orogeny on one part of a plate may be sufficient to cause folding in other parts of the same plate including uplift of a formerly passive margin on the opposite side of the continent from the orogeny. 119 cobbold et al. (2001, 2007) observed that post-rift uplift phases in brazil in the campanian, eocene and miocene are synchronous with the peruvian (90–75 ma), incaic (50–40 ma) and quechuan (25–0 ma) uplift phases in the andes. subsequent studies confirmed this result (cogné et al. (2012) and extended the synchroneity of all three events to the west african margin (japsen et al. 2012b; see fig. 70). the andean phases coincided with rapid convergence on the western margin of south america and uplift in the campanian coincided with a decline in spreading rate at the mid-atlantic ridge (torsvik et al. 2009). because the uplift phases in brazil and africa are common to the margins of two diverging plates, japsen et al. (2012b) suggested that the driving forces can transmit across the spreading axis, so their origin must therefore be in the asthenosphere, and that the common cause for both vertical movements and lateral changes in the motion of the plates is lateral resistance to plate motion. other authors have noted similar correspondence elsewhere. cloetingh et al. (1990) showed that rapid late neogene subsidence and sedimentation around the north atlantic were consistent with rapid changes of intraplate stress that most likely also gave rise to major changes in plate motions at that time. janssen et al. (1995) found a correlation between changes in plate motions and the evolution of rifted basins in africa. a phase of uplift and erosion of margins on both sides of the north atlantic (green & duddy 2010; japsen et al. 2010) at the eocene–oligocene transition (c. 35 ma) happened at the same time as a reorganisation of the spreading ridge there (gaina et al. 2009). these observations again indicate the presence of an influence that can transmit across a spreading ridge, and that uplift events of epcms correlate with changes in plate motion. the most likely source of these effects and of large stresses within the crust is basal traction from asthenospheric currents. alvarez (2010) argued that the continued collision of india with asia, after the descending slab had detached from the subducting continent, cannot be explained by slab pull and that ridge push alone is insufficient. thus some basal traction must also be involved. husson et al. (2012) argued that uplift of the andes can be explained by basal traction arising from an upwelling plume under southern africa pushing the south american plate against the subducting nazca plate. this model implies that the entire south american plate is in compression (cf. cobbold et al. 2007), and this compression may be sufficient to fold the continental crust and cause uplift of eastern brazil. horizontal, asthenospheric flow on its own does not produce uplift, unless the flowing material is anomalously hot (when the uplift is due to pratt isostasy). the horizontal flow simply carries the overlying plate with it (like a boat on a river). the flow moves past the base of the lithosphere, creating basal traction and compression, only when the plate meets resistance (the boat hits a rock). in contrast to the correspondence of events on either side of the south atlantic, the miocene events of regional uplift and exhumation around the north atlantic began at quite different times in west greenland and southern scandinavia. in greenland, the uplift of the ups began at c. 10 ma (section 5.5) whereas the onset of miocene uplift and exhumation around southern norway occurred in the earliest miocene at c. 23 ma (japsen et al. 2007a; rasmussen 2008, 2010), and this movement is reflected in a regional, base-miocene unconformity in the nw european margin (stoker et al. 2005). these observations indicate that the driving forces for the uplift of southern scandinavia in the earliest miocene are related to forces transmitted within the eurasian plate rather than in the asthenosphere, which is why they did not affect west greenland. none of the effects of mantle-driven uplift or compression are, however, long lasting. changes in mantle upand downwelling may take place in very short times, less than 1 million years (e.g. in the palaeogene of west greenland, see chapter 5 and in the faroe-shetland region of the uk north atlantic margin, see shawchampion et al. 2008). changes in stress necessary to produce third-order sequences take place on a time-scale of 3 million years or less (cloetingh & ziegler 2009 and references therein) and changes in the compressional stresses large enough to cause crustal-scale folding must also take place on timescales less than 10 million years (see section 8.1). 7.5 summary the shared characteristics of epcms (high topography with inland sloping plateaux cut by incised valleys and steeper oceanward decline) suggest that epcms have similar underlying causes. we thus reject the assertion by maupin et al. (2013, p. 20) that “each mountain chain has its own characteristics, making it unique and leading to different preferred mechanisms for their tectonic evolutions”. causes that require special conditions that apply to only one or a small number epcms are unlikely 120120 to provide satisfactory explanations for the formation of epcms in general. the numerous observations presented here indicate that epcms form long after rifting and break-up as a result of episodic burial, uplift and exhumation. as discussed in section 7.4, it seems likely that some epcms are, at least partially, uplifted by upwelling mantle diapirs and/or hot underlying asthenosphere derived from them. there is no evidence, however, for present-day upwelling mantle diapirs near e.g. the brazil or east australian epcms, so this explanation seems insufficient on its own to account for all epcms. some other effect seems to be necessary, and one likely explanation may be lithosphere-scale folding combined with flexed isostatic response to local erosion onshore and sediment loading offshore (cloetingh et al. 2008). the lithosphere-scale folding is caused by compression, itself derived either from orogenies elsewhere on a plate or from basal drag of the lithosphere by horizontal asthenospheric flow. the effects of these two sources of the forces driving vertical motion along epcms may explain why both peneplains and events of cooling/exhumation detected by afta are regional in extent and also explain why some uplift events are common to margins of divergent plates whereas some are not. 121 8. key issues concerning the development of epcms 8.1 steady-state or transient landscapes, ancient landscapes or young landscapes? that landscapes are in a steady state is a view that is held by many researchers modelling isostatic uplift of continental margins (see references in bishop 2007), and summerfield (2000) and others have cast doubt over the idea that landscapes have a memory indicating former base levels. the results discussed in chapters 5 and 6 question these views. even though it is clear that a landscape is all the time developing toward a steady state, there are reasons to believe that there is a considerable time lapse between cause and effect in the development of landscapes (cf. brunsden 1993), giving rise to what are now sometimes called transient landscapes (see bishop 2007). formation of extremely low-relief landscapes such as the sub-cambrian peneplain demonstrates that it is possible to produce a peneplain as the end result of erosion (section 3.1.1). a tectonic uplift event is needed to change from a situation of subsidence to re-expose the old landscape, and tilting will produce valley incision (section 3.1.5). isostatic compensation for erosion does not, however, mean that isostasy maintains the height above sea level of a landscape despite erosion. erosion always reduces the average height of a landscape, although isostacy causes the reduction to be less than the amount eroded; to reduce the mean elevation of a region by a certain amount, a much larger thickness of rock must be removed. assuming airy isostasy, a thickness of rock δt = δh(1ρc/ρm)-1 must be eroded to lower the height above sea level by δh, where ρc and ρm are the densities of material being eroded and mantle respectively. figure 73 shows that to denude a region to a low elevation in basement (i.e. to form a peneplain) a considerable thickness of rock must be removed, whereas to achieve the same in sedimentary rocks a much smaller removed thickness is required, because of the lower density of the eroded rocks. this may help to explain why in many areas, creation of peneplains in basement is associated with exhumation from palaeotemperatures >110°c, as recorded in afta data, whereas subsequent deposition of sedimentary cover and removal to form a new peneplain often involves removal of a smaller thickness of section, and lower palaeotemperatures. for example, the presently exposed basement of e.g. west greenland was metamorphosed at depths >20 km during the late proterozoic (c. 1600 ma), and that amount of upper crust had been removed prior to the ordovician (400 ma). in contrast, subsequent events have involved deposition and removal of much smaller thicknesses of phanerozoic cover. the history of the continental margin of north-east greenland illustrated in figs 74, 75 provides another illustration of these points. in the old studies of continental margins, the summit plateaux were thought to be graded to pre-uplift sea level and uplift was regarded to be of neogene age (e.g. reusch 1901; craft 1933; ahlmann 1941), while more recent studies were based on the idea that these surfaces were equivalent to the pre-break-up surface and thus of mesozoic age (e.g. ollier 1985). crosscutting relationships between peneplains (fig. 11) are important to determine, in order to draw conclusions on how far the reexposed peneplains extend and to ascertain where erosion has cut deeper than a re-exposed peneplain. the old ages sedimentary rock densities 7 ∆t – c ha ng e in c ru st th ick ne ss (k m ) 6 5 4 3 2 1 0 1700 1900 2100 2300 density of eroded rocks (kg/m3) 2500 2700 2900 basement densities fig. 73. plot of the amount of erosion (change in crustal thickness) needed to reduce the mean elevation of an uplifted area by 1 km vs. the density of the eroded rocks (assuming local airy isostasy and a mantle density ρm = 3300 kg/m3). erosion always leads to a reduction in elevation, but, because of isostasy, the reduction is less than the amount eroded. erosion and isostasy do thus not lead to a ‘steady state’. if basement rocks of density ρc = 2800 kg/m3 are eroded, 5.7 km of erosion must take place (i.e. the crust thins by this amount) for every km of elevation reduction. in contrast, removal of only c. 2.5 km of sedimentary cover (ρc = 2000 kg/m3) is needed to lower the elevation above sea level by 1 km. 122122 suggested in many morphological studies of peneplains of the gondwana continents (ollier 1982; twidale 1985, 2007) might be explained by extrapolating too far from covers without noting changes in inclination of the peneplains and/or interpreting such changes as caused by warping instead of deeper erosion in the basement and formation of a younger peneplain. the older idea of young (late palaeogene/neogene) ages of high-level peneplains (e.g. reusch 1901; ahlmann 1919, 1941) seems to be more realistic when tested against afta (chapters 5 and 6). japsen et al. (2006, 2012b) investigated the time available for the formation of peneplains in west greenland and north-east brazil and found that an interval of 20 million years is sufficient, after rejuvenation of relief during an initial uplift event (see fig. 70). japsen et al. (2006, 2012b) showed that in west greenland, the present-day relief formed after uplift that began in the late miocene (c. 10 ma; see fig. 54) whereas the relief in north-east brazil formed after uplift that began in the early – middle miocene (between 18 and 15 ma; fig. 68). it seems possible that in many places around the world, much of the present-day relief is not older than the neogene/late palaeogene (cf. thornbury 1969), unless it has been re-exposed from below a former cover. how far back in time the present landscape has a memory seems to depend on bedrock resistance, climate, and location (twidale 1976; brunsden 1993; bishop 2007). recently, egholm et al. (2013) drew renewed attention to a long-standing debate (also see e.g. baldwin et al. 2003) concerning the apparent enigma of the long-term preservation of kilometre-scale mountainous relief in ancient palaeozoic orogenic belts which are tectonically inactive today, citing examples including the caledonides of greenland and scandinavia and the lachlan orogen of south-east australia. in the light of the results from these areas presented here, this enigma can be simply resolved by acknowledging that the mountains in these classic top of sampled section must have been close to the surface when palaeogene volcanics were deposited 0 carboniferous permian trias. jurassic cretaceous cenozoic 10 20 30 40 50 60 te m pe ra tu re (° c ) 70 80 90 100 110 120 400 350 300 250 200 time (ma) 150 100 50 0 2 3 4 5 1 more realistic histories involve reheating by at least 50–70°c in the mid-cenozoic fig. 74. thermal history solutions for a vertical section of carboniferous sandstones on clavering ø, east greenland (c. 74°n; location shown in fig. 75), from johnson & gallagher (2000). also shown is a more realistic interpretation obtained by incorporating the presence of the palaeogene volcanics at the highest elevations above the section sampled by johnson & gallagher (2000). these constraints show that the uppermost sample must have been at near surface temperature in the palaeogene, implying a major degree of re-burial by additional volcanics and/or post-volcanic palaeogene sediments prior to the onset of cooling from the mid-cenozoic palaeothermal peak of about 70–80°c. the alternative cooling history indicated matches two episodes of cooling at c. 35 and 10 ma (thomson et al. 1999). each line represents the temperature history of an individual sample. in section (1), fission tracks are produced and annealed in the source region of the sediment prior to deposition; in section (2), maximum temperatures at 274 ma (mid-permian); in section (3), cooling at 206 ma (early jurassic); in section (4), poorly constrained cooling at 140 ma (early cretaceous); and in section (5), cenozoic reheating and maximum temperatures at 23 ma. modified after japsen et al. 2010. 123 epcm regions and others are the result of much more recent rejuvenation and uplift, unrelated to the ancient orogenies. as described in chapters 2 and 3, the form of landscapes should be regarded as a source of information on their development. this information can then be used to design more accurate modelling studies, rather than designing such studies to explain a particular view by choosing appropriate parameters. 8.2 landscape evolution and thermochronology: progressive emergence (continual cooling) vs. episodic burial and exhumation (heating and cooling) as discussed in preceding chapters, a common assumption in thermochronological studies of epcms is that the rocks comprising the uplifted areas have cooled monotonically. this presumably reflects the widespread notion that rocks in such regions reach the present-day surface through a long history of relatively slow denudation. this assumption is difficult to reconcile with the observations from west greenland presented in chapter 5, where the exposed geology documents a series of episodes of burial and exhumation. as discussed in chapter 6, results from other epcms, including brazil, south africa and south-east australia can be interpreted in terms of a similar episodic style of evolution. in addition, sla in southern scandinavia provides clear evidence of episodic evolution (chapter 3). the concept of monotonic cooling appears to derive from classical models of landscape development. major landscape forms such as low-relief erosion surfaces in stepped sequences led to the idea of continuous uplift interrupted by periods of quiescence, and the formation of vast peneplains (davis 1899) or pediplains (king 1951). king (1962, 1967, 1983) acknowledged the former existence of covers, e.g. a former cretaceous cover of the brazilian margin, and noted that there had been periods of subsidence and burial. yet, working with landforms, he mainly discussed the uplift events. in europe the idea of a continuous lowered base level since the cretaceous was used by baulig (1935) to argue that this was the cause of erosional steps seen in the topography (the ‘eustatic theory’ referred to by chorley 1965; also see chapter 2) as by then the continents were thought to be stable. thus the highest surfaces were considered the oldest; it was thought that either uplift had taken place or that sea level had been lowered. in general, the old studies did not consider re-burial as an important factor in landscape development, because in those days there was no way of documenting whether thick sections of rock might have been removed. thus landscapes were analysed with the assumption of a monotonic, though intermittent rise relative to sea level. we know now that eustatic sea-levels have never varied by more than a few hundred metres (miller et al. 2005), and evidence has now accumulated that disproves the notion of continental stability (e.g. moucha et al. 2008). despite these changes in understanding, the notion of continual emergence, which was originally devised as an essential component of the eustatic explanation, has remained from times that pre-date the acceptance of a more mobile crust dominated by mantle-driven processes in a plate-tectonics environment (see chapter 7; e.g. steinberger 2007; roberts & white 2010; jones et al. 2012). re-exposed surfaces of different age (ambrose 1964; twidale 1985) are indicators of former covers. yet such surfaces have not been commonly used in the past to reconstruct landscape development. on land, such covers are just remnants and originally must have been both thicker and extended farther, even outside the present distribution of re-exposed relief, which is commonly cut off by a younger peneplain (figs 11, 14). such covers seem to have been the norm. covers of different age directly on precambrian basement in fennoscandia indicate that the major erosional phase of basement rocks occurred before the cambrian (lidmar-bergström 1997). major periods of covering in scandinavia were then the palaeozoic, the late mesozoic, and probably the eocene, while most of the mesozoic, the paleocene and post-eocene were periods of uplift and erosion in the basement (lidmarbergström 1995, 1996). lidmar-bergström et al. (2013) thus showed that the relation between relief in basement and cover rocks of different age directly on basement in scandinavia provides important information about the phanerozoic uplift and subsidence history of the region. in the approach favoured by some geomorphologists (section 2.4), re-exposed relief is often assumed to be evidence of removal of just a thin cover, and this assumption may lead to difficulties in accepting the results of apatite fission-track data (gunnell 2000; peulvast et al. 2008; also see section 6.4). where sedimentary outliers are present on basement rocks, they provide an important constraint, defining times when the basement surface was previously at or very close to the land surface. in such circumstances, 124124 representative sample and sample number aft sample, pedersen et al. (2012) 20 20 30 28 vr c. 0.5% vr c. 0.5% vr c. 0.5% palaeogene sediment palaeogene basalt middle jurassic – cretaceous sediment upper carboniferous – permian sediment caledonian basement ice pre-caledonian rocks sedimentary outlier 100 km store koldewey kulhøj depotnæsset germania land kuhn ø clavering ø hochstetter forland holm land 12°w 12°w 20°w 20°w 74°n 76°n 78°n 80°n 0 ord si dev carb pe tr jur cret cen 1 2 3 4 5 0100200300 time (ma) 400500 group 1 sog20 group 2 sog30 group 3 sog28 exhumation and cooling histories from pedersen et al. (2012) 1 3 >3 k m d isc re pa nc y be tw ee n m od el a nd d at a 5 d ep th (k m ) a 2 4 2 4 6 bb c d 2 4 1: caledonian metamorphism; 700–800ºc 2: exhumed to surface by late carboniferous 3: buried by 2 km of u. carboniferous to pre-middle jurassic sediment 4: exhumed to surface by middle jurassic 5: buried by 1–2 km to reach c. 80ºc maximum post-jur. palaeotemp. 6: exhumed to surface at present day rocks exhumed to the surface 125 interpretations involving monotonic cooling histories will produce meaningless results. this may seem obvious, but a number of published examples have failed to take the presence of cover rocks into account (johnson & gallagher 2000; persano et al. 2006; pedersen et al. 2013, as illustrated in figs 74–77. clearly, thermal histories involving continuous cooling are inappropriate in areas where cover rocks are preserved. integration of the constraints provided by cover rocks is essential in order to obtain realistic thermal histories. 8.3. epcms: permanent highs or the result of late uplifts? the development of the relief of the natal monocline after formation of the african surface (below the drakensberg escarpment) resulted in lower steps in the landscape (several generations of peneplains and valleys; fig. 2). the only explanation presented so far for these stepped surfaces and incised valleys is grading to lowered base levels following neogene uplift events (king 1972, 1983; partridge & maud 1987). the presence of both several high surfaces and lower valley steps is neglected in the prevailing approach to such matters as discussed in section 2.4. ollier & marker (1985), for example, who studied the great escarpment of southern africa, only identified one palaeoplain that they argued to be of mesozoic age. they assumed that the margin had remained high since break-up, and suggested two models for its development, marginal downwarp and rift-shoulder uplift (fig. 57). these ideas were adopted in the models created and tested by e.g. brown et al. (2000) and persano et al. (2002). these authors adopted the new methods of thermochronology and focussed on the classic areas of lester king in southern africa and of cliff ollier in eastern australia. their models were based on three assumptions formulated by gilchrist & summerfield (1994). first, that the present topography has an origin related to continental break-up. second, that the marginal upwarps of mature passive margins cannot be explained by the dynamic effect of rifting as the axis of maximum uplift is now located 100 km or more inland and therefore they regard isostatic response as an important factor (cf. king 1956b). third, that the new continental margins were originally at high elevation at break-up due to either rift-related surface uplift or pre-existing residual high terrain. their idea, which many have accepted, is that the continental margins reached their present elevations before break-up of gondwana and have remained elevated ever since. the idea of an originally high margin is nowadays often taken for granted (e.g. campanile et al. 2008). in contrast, early observations (section 8.1) of the relationships between high-level peneplains and incised valleys gave a different explanation of epcm development as a mainly neogene story. this description is consistent with our results from west greenland and elsewhere (chapters 5 and 6). as shown in these chapters, there is now a strong body of evidence to show that passive margins have moved up and down and that their recent elevation is coupled to and maintained by geodynamic processes hitherto not completely understood. facing page: fig. 75. conflicting interpretations of the post-caledonian development in north-east greenland illustrating the importance of stratigraphic constraints in the interpretation of thermochronology data. a: geology of north-east greenland and location of apatite fission-track (aft) samples analysed by pedersen et al. (2012). outliers of upper carboniferous – permian and jurassic–cretaceous sedimentary rocks (triangles coloured according to the legend) occur in patches between 75° and 80°n, while these sedimentary units are more extensive farther south and farther north. the presence of these sedimentary rocks demonstrates that the underlying basement was at the surface prior to their deposition. grey circles: samples sog20, 28, 30 that pedersen et al. (2012) chose as representative of their three groups of samples. b: outcrop of middle jurassic coal located at triangle 4 on the map. c: upper carboniferous deposit located at triangle 2 on the map; fossil and reconstruction of lepidodendron (piasecki et al. 1994). d: exhumation paths (red, blue and green) based on modelling by pedersen et al. (2012) of apatite fission-track data in three representative samples chosen by these authors. superimposed curves (yellow) illustrate the geological constraints on the burial and exhumation history of basement rocks around germania land. (1) caledonian basement with eclogitic inclusions reflects high-pressure metamorphism of devonian age (700–800°c, 410–390 ma; gilotti et al. 2008). (2) the presence of upper carboniferous sediment overlying the basement shows that the basement was exhumed to the surface by the late carboniferous (piasecki et al. 1994). (3) maturity of sporomorphs in these deposits indicates that they were buried below a cover, 1.5–2 km thick (piasecki et al. 1994). (4) the presence of middle jurassic sediment overlying the basement shows that the upper carboniferous – middle jurassic cover was partly removed by the middle jurassic (e.g. bojesen-koefoed et al. 2012). (5) vitrinite reflectance (vr) values of c. 0.5% for the jurassic deposits show that they were buried below a sedimentary cover, 1–2 km thick (bojesen-koefoed et al. 2012). (6) today the caledonian basement is exposed at the surface over most of the region. modified from japsen et al. (2013b). 126126 20 0 40 60 80 100 120 te m pe ra tu re (° c ) 20 0 40 60 80 100 120 te m pe ra tu re (° c ) 100 0200300400 100 0200300400 jurassic cretaceous cenozoictriassicpermiancarbonif.dev. a. episodic cooling and heating history (o’sullivan et al. 1995) b. monotonic cooling history (persano et al. 2006) jurassic cretaceous cenozoictriassicpermiancarbonif.dev. time (ma) time (ma) geological constraints: 1: carboniferous emplacement of bathurst batholith (310 ma) 2: rapid post-emplacement cooling of batholith 3: granite now at outcrop exhumed to surface by mid-permian 4: re-burial of granite as a result of deposition of permian to triassic sydney basin sequence original interpretation of aft data (o’sullivan et al. 1995): 5: cooling at 90±10 ma from maximum temperatures increasing from 80–90°c in west (blue) to 90–100°c (green) in central region to 100–120°c in the east (red) 6: ‘moderately rapid’ late cretaceous cooling interpretation of aft and apatite (u-th)/he data without geological constraints (persano et al. 2006): 7: granite outcrop samples cooling from temperatures >100°c in late permian to early triassic, when the permian–triassic sedimentary sequence (4) was being deposited onset of cretaceous cooling (o’sullivan et al. 1995) rocks exhumed to the surface 4 2 1 1 5 6 7 4 2 3 3 fig. 76. two conflicting interpretations of the post-carboniferous development in south-east australia, illustrating the importance of stratigraphic constraints in the interpretation of thermochronology data. a: thermal history interpretation of granite samples from the bathurst batholith, west of sydney, new south wales, from o’sullivan et al. (1995), together with constraints provided by geological evidence, viz: (1) emplacement of the batholith at c. 310 ma (facer 1978); (2) rapid post-emplacement cooling as demonstrated by (3) granite now at outcrop (and sampled by o’sullivan et al. 1995) overlain by permian sedimentary units of the sydney basin sequence; (4) deposition of permian to triassic sedimentary units of the sydney basin sequence above present-day granite outcrops. o’sullivan et al. (1995) interpreted apatite fission-track (aft) data in samples of both granite and overlying permian sedimentary remnants as defining heating to early cretaceous temperatures (5) from 80–90°c in the west to 90–100°c at central locations and 110–120°c in the easternmost outcrops where the granite disappears beneath the sedimentary cover, followed by rapid cooling (6) which began in the interval 90±10 ma. b: persano et al. (2006) integrated apatite (u-th)/ he ages from samples of outcropping granite with the apatite fission-track data of o’sullivan et al. (1995) and interpreted the results in terms of slow cooling, with samples in the west and centre of the study region cooling through c. 100°c in permian times (7), while samples farther to the east only cooled below c. 100°c in the early cretaceous. the geological constraints (1–4) show that the granite now at outcrop was close to the surface during the late permian and being buried by the sydney basin sequence. persano et al. (2006) did not take the geological constraints into account, resulting in unrealistic conclusions (see green et al. 2007; brown 2007; gibson 2007 for further details). 127 8.4 the meaning of ‘escarpment’ the notion of a ‘great escarpment’ is often confusing in studies of passive-margin development. it sometimes denotes the whole oceanward slope of the epcm, sometimes the deeply incised valleys below the base of the high-level landscapes and sometimes a structural feature such as the drakensberg escarpment in south africa (fig. 2). we use the word escarpment in its broadest sense, as a steep slope. conspicuous, high escarpments are often made up of resistant rocks; e.g. the drakensberg escarpment (moore & blenkinsop 2006), the escarpment around beaufort west in south africa dominated by karoo sills and the escarpment along the south-eastern australian coast south of sydney in triassic sandstones. the notion ‘great escarpment’ was used by ollier (1982) for the highly irregular backwall caused by the incised valleys, which abruptly cut into the elevated and arched plateau surface of eastern australia. there are valleys incised into the african surface in southern africa of the same nature as those making up the so called great escarpment in eastern australia, but these valleys are incised into the african surface and thus start below the drakensberg escarpment which forms part of great escarpment defined by ollier & marker (1985) in southern africa. this difference in the definition of great escarpment in the two areas is not noted in recent literature on modelling of epcms, and the existence of a stepped landscape on the oceanward side (fig. 2) is ignored. in the king (1956b, 1972, 1983) scheme, the drakensberg escarpment is just one, though prominent, of a series of steps. ollier (1985), however, did not acknowledge the relevance of a stepped topography and only discussed the case of one high-level plain and one great escarpment forming the general coastal slope. it is this description that most modern modellers have used even though it is often not applicable. the review of the landscapes at different epcms presented here suggests that while different epcms share many characteristics (e.g. the presence of high-level, lowrelief surfaces cut by deeply incised valleys), many differences in detail exist, which provide further information on the development of specific margins. high-level surfaces are often present at several elevations (different for each margin), but often there is a lowest major peneplain below which rivers have cut deep gorges down to the coastal plain. at some locations the valleys have widened to give rise to an escarpment of high sinuosity, as seen in eastern australia (fig. 3), probably due to deeply weathered bedrock. in glaciated areas, incised valleys have often been elongated backwards and form long fjords. at some locations resistant rocks have formed a wall, a real escarpment, but this wall can be just one step of a margin with several steps. thus climate and rock structure influence escarpment appearance. for conclusions on the tectonic development of epcms, it is the high-level, lowrelief surfaces that provide the most important insights, whereas the exact appearance of the escarpment is of minor importance. 8.5 down-wearing, back-wearing and scale there is an important difference between the processes acting on high-level palaeoplains and those that shape deeply incised valleys. high-level palaeoplains are characterised by slow down-wearing processes (the general lowering of a land surface by denudation), while there is an intensified weathering along incised valleys, due to reactivation of groundwater flow, accelerated slope processes and backward retreat at valley heads (ollier 1982; thomas 1994). some researchers regard deep weathering/stripping of saprolite as the major agent in landscape formation. they refer to these processes as downwearing and consider base level to be of no importance (thomas 1994; phillips 2005). ‘etch surface’ is then a label for the entire landscape, which is regarded as being in a steady state as denudation (weathering and erosion) keeps pace with uplift, but maybe with a change of saprolite thickness due to climatic changes (thomas 1989). in sla, base level is instead regarded as decisive also for etch-surface formation, as erosion of the landscape is limited by this base level. the etching (deep weathering) is a secondary feature, dependent on climate. japsen et al. (2005, 2006, 2009) showed that in west greenland the uplifted ups formed during oligocene– miocene times. over the comparatively short time span until today, a slow down-wearing process of the original peneplain, possibly after stripping of any unconsolidated cover, can be regarded as negligible (cf. phillips et al. 2006). thus the peneplain level can be used for estimation of rock uplift since peneplain formation. scarp retreat is the classical back-wearing concept (king 1962, 1967; ahnert 1982, 1998). in recent papers, the concept of down-wearing has been used to describe long-term river incision and valley widening which results in major erosion of continental margins (gallagher et al. 1998; persano et al. 2002), thus giving down128128 wearing the same meaning as back-wearing in classical studies. the idea of a pinned (stationary) divide along the line of the present escarpment of epcms was suggested by gallagher et al. (1998), meaning that its position was due to different rates of denudation between the divide (slow denudation), the inland areas (higher denudation), and the coastal area (highest denudation; fig. 57c). yet it is evident from profiles by king (1972, 1983) and a profile by van der beek et al. (2002) that the drakensberg escarpment is far northwest of the uplift axis of the natal monocline (section 8.4). erosion has caused considerable retreat of the escarpment and mainly across less resistant rocks. moore & blenkinsop (2006, p. 599) noted that the drakensberg escarpment is not restricted to the ocean-ward side and they demonstrated that this feature is due to headward retreat. they write: “headward retreat of large waterfalls on the major rivers over considerable distances (10–100 km) provides clear field evidence that scarps formed by resistant lithologies will not invariably degrade as a result of the existence of an inland drainage divide, as has been argued on the basis of surface process modelling.” that gorge-head ? 0 20 40 60 80 100 120 te m pe ra tu re (° c ) time (ma) 0100200300400 palaeothermal constraints from afta sample rocks exhumed to the surface regional constraints on the onset of cooling jurassic cretaceous cenozoictriassicperm.carbonif.dev.b 2 4 5 6 7 dwyka tillites (karoo) basement (late carboniferous glacial landscape) etendeka volcanics (early cretaceous) dwyka tillites (karoo) basement (late carboniferous glacial landscape) a 5 2 8 1 3 3 1: granite exhumed during late carboniferous 2: formation of glacial landscape 3: deposition of karoo cover 4: basement reheated to >110°c below karoo cover; cooling began c. 200 ma 5: removal of most of the karoo cover and deposition of etendeka basalts 6: basement and younger cover reheated to c. 90°c from which cooling began at c. 100 ma 7: later episode of cooling 8: exposure of basement and remaining cover at the surface today fig. 77. stratigraphic constraints on the thermal history interpretation of afta data from a basement outcrop in north-west namibia (18.184°s, 12.762°e). a: late carboniferous – early permian glacial landscape, as evidenced by the roche moutonnée on which the geologists are standing, which is actively being exhumed from beneath a former cover of the late carboniferous – early permian dwyka group of the karoo supergroup. this is overlain, in turn, by basalts of the lower cretaceous etendeka volcanics, comprising the distant peak. b: extraction of thermal history information from afta data in apatites from the granite of the roche moutonnée reveals three episodes of cooling, as shown by the coloured box outlines (4, 6 and 7). afta data from samples across the region provide constraints on the onset of four episodes of cooling/exhumation, as shown by the vertical, coloured bands. (1) samples farther inland preserve evidence of early cooling, representing exhumation of the granite to the surface. (2) geological evidence shows that the granite had been exhumed to the surface, and was therefore at low temperature, prior to deposition of the karoo supergroup. (3) the afta data show that the rock was then reheated to >110°c, implying deposition of a thick karoo cover, of which the surviving dwyka sediments are just a remnant. (4) afta data show that the granite sample cooled below 110°c at around 180 ma. (5) the presence of the lower cretaceous etendeka basalt only c. 100 m above the basement outcrop shows that by c. 132 ma the sample must have cooled to near-surface temperatures, implying that the former karoo cover had been largely removed by this time. (6) afta data define another phase of cooling from c. 90°c which began at c. 100 ma (based on regional data). this implies re-burial by some thickness of lower cretaceous section. (7) a later episode of cooling is also defined from the afta data. (8) today the granite of the roche moutonnée is exposed at the surface. integration of afta with stratigraphic constraints defines the episodic nature of the thermal history. interpretation of the afta data in terms of monotonic cooling clearly conflicts with geological evidence at this location. note that if erosion had been more pronounced, and the cover rocks removed, as over a large area of the surrounding country, no evidence would be preserved of this episodic heating/cooling history 129 migration is a most important process as a bedrock incision mechanism is also demonstrated in two papers from the east australian margin (seidl et al. 1996; nott et al. 1996). in their models of development of epcms, gallagher et al. (1998) and persano et al. (2002, 2005) present new ideas on how fluvial landscapes develop by river incision and valley widening along a pinned divide without headward retreat. their denial of headward retreat also seems to be the reason why the fluvial process is looked upon as down-wearing in these models, although they accept the process of valley widening which is part of a general scarp retreat. these are severe inconsistencies. in the long time span needed for formation of a full peneplain, the question of down-wearing contra back-wearing is not an important issue, compared to what is seen from the thermochronology constrained by sedimentary covers. no significant topography is left. this means that the present difference in process rates between the high plains and the incised valleys, is a relatively new feature. there is thus a scale difference between those processes causing a slow down-wearing of high-level plains and the rapid processes causing back-wearing by river incision and valley widening along epcms. 8.6 are epigene surfaces sometimes re-exposed? the discussion in chapter 2 distinguishes between reexposed and epigene peneplains; the former we identify at the contact with the cover, in the case of the latter, no cover is present. however, results discussed in chapters 5 and 6 suggest that the apparently epigene peneplains, the ups in west greenland and the higher surface in northeast brazil, have had covers (note that a palaeogene cover is preserved on some remnants of the higher surface in brazil; section 6.4.4). epigene peneplains are generally preserved after uplift because of their location far from incising rivers that destroy the peneplains by backwearing. how is it possible that the suggested oligo–miocene covers in west greenland and north-east brazil can be stripped over large areas and the underlying surface at the same time is totally preserved? difference in resistance to downwearing between a relatively soft cover and a resistant base provides a possible answer. the tilted ups of west greenland has almost no residual relief and might represent a re-exposed unconformity. the reason why the erosion surface exposed on westernmost nuussuaq is so well preserved could therefore be that it has been protected by the miocene cover which is preserved immediately offshore farther west (figs 41, 42a). in north-east brazil, the palaeogene serra do martins formation covers parts of the palaeogene higher surface on the borborema plateau but not on planalto da conquista (see section 6.4.4). this indicates that the surface is in fact a re-exposed unconformity. if the, by definition, epigene surfaces are sometimes reexposed from temporary unconsolidated, easily eroded covers, the label epigene peneplain just means that no covers are found on it. 8.7 episodic development of epcms in preceding chapters, evidence has been presented to show that a number of epcms have been uplifted significantly later than the time of rifting and continental break-up, and often in multiple episodes. japsen et al. (2012a) have pointed out that unconformities offshore are commonly extensions of the erosion surfaces present at high level in epcms. these unconformities commonly truncate earlier sedimentary successions that generally dip oceanward (fig. 6), showing that the sedimentary units once extended across the onshore region. these structural relations thus indicate episodic, positive and negative post-rift vertical movements at epcms. one simple explanation why many workers have assumed epcms to be permanent uplifts that are somehow related to rifting and break-up, may be that many epcms appear to be located above thicker crust/lithosphere in close juxtaposition to thinner crust/lithosphere; i.e. along formerly active rifts (japsen et al. 2006, 2012a; osmundsen & redfield 2011). this condition occurs landward of where continental crust starts to thin as rift basins towards oceanic crust (for example in brazil, south-east australia and scandinavia). but it also occurs where thicker cratonic crust adjoins thinner extended continental crust within a continent, such as in southern sweden where 40 km thick crust under the south swedish dome is juxtaposed to 30 km thick crust on the south-western side of the sorgenfrei–tornquist zone (japsen et al. 2012a, fig. 8). redfield & osmundsen (2013) also suggested that one single uplift event cannot have created scandinavia’s stepped topography. they found the evidence to be more consistent with a series of uplifts occurring episodically throughout the cenozoic, each followed by a 130130 period of stability during which erosional surfaces were incised to a relatively low base level. in their view, the progressive and episodic tilting of the norwegian hinterland caused slope-dependent landforms to become rejuvenated, reinforcing or overprinting the preceding landscape until a new equilibrium became established. redfield & osmundsen (2013) did not consider the possibility that the norwegian margin has been buried since late jurassic rifting. there is, however, evidence that sediment accumulated over the present onshore areas in jurassic times and possibly later; vr values for erratic blocks from a jurassic basin in a mid-norwegian fjord suggest exhumation of the basin from below a cover of c. 2 km (sommeruga & bøe 2002). this observation is in agreement with the presence of re-exposed sub-mesozoic relief on the mid-norwegian coast (fig. 7). the episodic nature of vertical movements at epcms presented here helps to clarify various issues that have long been debated. for example, in southern africa burke and gunnell (2008), tinker et al. (2008a) and paton (2012) have all highlighted an apparent conflict between evidence of cretaceous ‘uplift’ defined from thermochronology and evidence for cenozoic ‘uplift’ provided by more conventional methods, as discussed in chapter 6. a period of cretaceous exhumation (leading to formation of low-relief surfaces across much of southern africa) followed by cenozoic uplift with only limited accompanying or subsequent erosion (taking the surfaces to their present elevations) provides a plausible explanation of the modern-day topography, and there need be no conflict between the rival interpretations (see section 6.2). this discussion highlights occasional confusion in the literature in understanding the nature of events revealed by low-temperature thermochronology, and that a minimum of two post-rift uplift events are needed to shape the present-day relief with elevated plateaux cut by deep valleys. in summary, the development of epcms involves a series of positive and negative post-rift, vertical movements, in which a considerable amount of rock is both deposited and removed during the period subsequent to rifting and break-up. the episodic development described here raises the question of the origin of the sedimentary cover required to bury a region by kilometre-scale thicknesses of cover. given the nature of events described, a likely explanation is that while some areas are subsiding and accumulating sedimentary cover as a result of negative movements, other areas are being uplifted and eroded due to positive movements. identification of such paired events remains a challenge for future definition of the nature of the events responsible for the episodic evolution of epcms. a similar question regards the ultimate destination of the rocks eroded in such events. several attempts have been made to compare the sediment budget in offshore basins with the denudation history of onshore margins (e.g. campanile et al. 2008; tinker et al. 2008b; rouby et al. 2009; anell et al. 2010). however, such analyses do not take into account the possibility that the offshore sedimentary basins may have themselves been eroded. in many of the cases discussed here, offshore basins have undergone significant amounts of denudation, although the resulting unconformities show little or no angular discordance, and are interpreted as representing periods of non-deposition. in addition, it is sometimes difficult to determine the ultimate resting place of eroded sediments, which can be transported over wide distances. ager (1973) noted that at any one time, most of the sediment within a basin is moving laterally through the basin, rather than being retained. this being so, the question of mass balance between eroded sediment from a margin and thicknesses of sedimentary rocks in an adjacent basin seems far from simple. 8.8 observation vs. theory in understanding the evolution of epcms stratigraphic landscape analysis is based on observations of landscapes and geological information. when integrated with information from thermochronology, the combined approach has led to a description of the evolution of epcms in terms of episodic burial and exhumation (subsidence and uplift) which is fundamentally different to accepted ideas of epcm evolution, as illustrated in preceding chapters. yet this new description continues to draw criticism from many quarters (nielsen et al. 2009; redfield 2010; pedersen et al. 2012, 2013), largely on the grounds that no mechanism is available by which this style of development can be explained. in several areas of geology over the last 150 years, conclusions drawn from basic observations have been denied because theory ‘proves’ that those conclusions cannot be valid. examples include the occurrence of past ice ages, deep geological time, and continental drift. micro landforms, such as striae, puzzled researchers for many years, but their observations demanded an explanation, which was ultimately obtained with the theory of the ice age. the value of 100 million years (later 20 million 131 years, stacey 2000) for the age of the earth derived by the future lord kelvin (thomson 1863) for many years constituted a major problem for acceptance of natural selection (england et al. 2007). macro landforms such as the shape of the continents gave the idea of continental drift. the concept was supported by fossils on both sides of the atlantic which belonged to the same type of environments while nearby fossil communities were different, as well as numerous similarities in the geology of the two continents across the ocean. as such the process was clear to southern hemisphere geologists (e.g. du toit 1937). but geophysicists (mainly located in the northern hemisphere) argued that continental drift was physically impossible. such attitudes provided a major obstacle to understanding, instead of initiating a search for the underlying processes. acceptance of ideas for which evidence had become overwhelming was delayed for many years, until ultimately the underlying mechanism became clear. in the case of continental drift, it is ironic that the mechanism was not discovered as a result of a careful programme of investigation in an attempt to discover the mechanism, but only through a serendipitous combination of circumstances (allègre 1988). similar comments apply to the recognition of the role of radioactivity in providing a longer geological timesecale and revealing the depth of geological time. all the above examples are characterised by a failure to recognise that accepted theories no longer provided an explanation of available data, and that new ideas were required. in the case of the evolution of epcms, evidence has been presented in previous chapters which leads to the conclusion that prevailing notions are no longer acceptable, and a paradigm shift is required. yet studies involving monotonic cooling and continuous denudation continue to appear, the simplistic models shown in fig. 57 continue to be popular (bishop 2007; campanile et al. 2008; burke & gunnell 2008), and the idea of removal of a significant thickness of cover continues to be regarded as unrealistic (e.g. peulvast et al. 2008). numerical modelling techniques have advanced in recent years. the notion of the continuous elevation of epcms, however, remains as a foundation (e.g. sacek et al. 2012), despite abundant evidence to the contrary (e.g. cobbold et al. 2001). in an influential paper on ‘geomorphology as a science: the role of theory’ rhoads & thorn (1993, p. 287) define science as an activity that “seeks to discover knowledge through a two-stage activity involving the creation and justification of ideas (theory)”. they go on to argue that “the primacy of observations in science is a myth and that all observations are theory-laden in the sense that the act of observation inherently involves interpretation and classification”. while this is true, it is also clear that if all observations are interpreted purely in terms of accepted theories, then no new discoveries will ever be made, so the mind must remain open to new concepts when existing models fail. in the examples cited above involving ice ages, the age of the earth and continental drift, adherence to existing theories constituted an obstacle to progress. a wealth of information shows that traditional theories for the development of epcms are no longer satisfactory. it is a first-order observation that high-level regions of low relief characterise epcms. such regions of low relief, cutting across rocks of different age and resistance, can only be created by erosion to base level, and since epcms are formed adjacent to oceans, in the absence of resistant lithologies over wide areas, the most likely base level is sea level. as shown here, epcm landscapes demonstrate an episodic evolution involving a series of positive and negative vertical movements. however, instead of seeking models and theories that honour these observations, their significance is denied because accepted theories cannot explain them. this argument therefore seems to be back to front. new theories are required which attempt to explain the wealth of observations pointing to the episodic upward and downward movements of epcms, as well as the significant amounts of rock that are deposited and removed following continental break-up before the present-day epcm landscapes were developed. 132132 9. summary and conclusions the elevated, passive continental margin (epcm) of west greenland developed through a series of vertical movements (both positive and negative) following continental break-up (chapter 5). the topography of the west greenland margin shares many characteristics with other epcms (elevated plateaux 1–2 km a.s.l. with deeply incised valleys, gently declining inland and a much steeper oceanward slope terminating at a coastal plain) suggesting a common overall style of development. however, the still prevailing view of epcm development is that they retain their elevation and characteristic landforms from breakup or even earlier. this contrasts starkly with the history of episodic subsidence/burial and uplift/denudation described for the west greenland margin in chapter 5. previous thermochronological studies favouring the long-term elevation of margins have been based on assumptions of progressive uplift/denudation and continual cooling. however, a review of geomorphological and thermochronological studies from other epcms (chapter 6) shows that these assumptions often conflict with geological evidence, and histories involving repeated episodes of burial and exhumation (heating and cooling), similar to that defined in west greenland, are more realistic. a wealth of evidence supports the conclusion that the elevated, low-relief regions that characterise epcms are uplifted peneplains originally graded to sea-level in the adjacent, opening ocean. the peneplains were subsequently uplifted to their present levels by movements that took place long after rifting. in many cases the peneplains were preserved prior to uplift by a sedimentary cover, which was subsequently removed during uplift. following uplift the peneplains were further modified by valley incision and valley widening, sometimes in multiple episodes, leading to the development of additional peneplains at lower levels. this results in characteristic, stepped high surfaces and a low generation of incised valleys, sometimes coalescing to an escarpment. however, the description of the oceanward part of a margin in terms of a single great escarpment is over-simplified and most margins are more complex (chapter 8). overall, the typical epcm landscapes should be viewed as reflecting the integrated effects of their post-rift development, and are not related to the rifting process. while the shared characteristics of epcms suggest that they are likely to have formed as a result of essentially similar processes, the form of each epcm is different in detail, and the differences between margins can be used to provide conclusions regarding the specific development of each, using sla in concert with geological evidence and thermochronology. while the nature of the controlling process(es) remains unclear, the regional extent of the vertical movements documented here suggests a plate-scale control, possibly related to compressive stresses resulting from distal orogenies, basal drag of the lithosphere by horizontal asthenospheric flow or from changes in plate motion (chapter 7). such explanations may explain why some episodes are common to margins of divergent plates whereas some are not. the results presented here provide a consistent body of evidence to demonstrate that epcms have moved vertically, both downwards (subsidence/burial) and upwards (uplift/exhumation) long after continental breakup. definition of these movements in time and space is essential to a full understanding of the nature of the underlying processes involved in epcm development. more fundamentally, identifying common aspects between different epcms has the potential to allow investigation of the properties of the lithosphere and mantle which control these vertical movements of rifted margins, long after continental break-up. acknowledgements we appreciate the support from the bureau of minerals and petroleum (government of greenland), the carlsberg foundation, the danish council for independent research/natural sciences, stockholm university and the swedish research council that funded our studies in greenland. we are grateful for the interest and active support we received in the early stages of our work in greenland from many geologists at geus and the university of copenhagen, in particular from the late t.c.r. pulvertaft. i. duddy and s. holford are thanked for useful discussions that helped to shape the final form of this paper. paul a.m. andriessen, sierd cloetingh and piotr migoń are thanked for constructive reviews. 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related to the evolution of elevated passive continental margins (epcms) 2.1 peneplains as key data for understanding denudation, subsidence and uplift of epcms 2.2 historical review of the identification and use of peneplains in landscape analysis 2.3 criticism of landscape analysis and the idea of steady state in geomorphology 2.4 different approaches to geomorphological studies of passive margins 3. stratigraphic landscape analysis: a new approach for extracting histories of denudation, subsidence and uplift 3.1 observations and techniques of sla 3.1.1 the sub-cambrian peneplain 3.1.2 the south swedish dome 3.1.3 cross-cutting relationships and their implications 3.1.4 peneplains as unconformities in other areas and their crosscutting relationships 3.1.5 sea level, a major base level 3.1.6 other implications of re-exposed peneplains 3.1.7 peneplain formation discussion 3.2 terminology 3.2.1 landscape, relief and topography 3.2.2 peneplanation or pediplanation 3.2.3 pediplains (rock-cut plains), pediments, etched surfaces (hilly relief), inselberg plains and etchplains 3.2.4 classification of peneplains 3.2.5 high-level benchlands (stepped peneplains), incised valleys and the great escarpment 3.3 palaeoplains, preservation, destruction and age 3.3.1 effects of glaciation 3.4 combining sla and thermochronology 4. low-temperature thermochronology 4.1 apatite fission-track methods 4.1.1 historical background 4.1.2 thermal response of fission tracks in apatite 4.1.3 variation in annealing kinetics between different apatite species 4.1.4 extracting thermal history information from apatite fission-track data 4.1.5 monotonic cooling vs. episodic heating and cooling 4.1.6 the meaning of a fission-track age 4.1.7 ‘boomerang plots’ 4.1.8 uplift rates from apatite fission-track age profiles 4.1.9 long term residence in the partial annealing zone vs. heating and cooling 4.1.10 limitations of apatite fission-track methods 4.1.11 alternative views of fission-track annealing 4.1.12 summary 4.2 apatite (u-th)/he dating 4.2.1 historical background 4.2.2 early success 4.2.3 evidence of greater complexity 4.2.4 he-closure temperatures and ‘cooling ages’ 4.3 converting thermal history information to denudation history 4.3.1 denudation histories 4.3.2 palaeogeothermal gradients and removed section 4.3.3 elevated heat flow on continental margins 4.4 other methods for constraining removed section and denudation histories 4.4.1 introduction 4.4.2 vitrinite reflectance (vr) 4.4.3 zircon fission-track analysis (zfta) and zircon (u-th)/he dating 4.4.4 estimating palaeoburial using sonic velocity data 4.5 constraints from basic geological data and stratigraphic landscape analysis 5. the west greenland margin; a consistent synthesis of geological data, stratigraphic landscape analysis and low-temperature thermochronology 5.1 an integrated approach 5.2 geological background: the nuussuaq basin and the offshore record 5.2.1 onshore exposures of the nuussuaq basin 5.2.2 the geological record off southern west greenland 5.3 results of sla in west greenland 5.4 use of palaeothermal methods to define an absolute chronology of uplift and exhumation events 5.5 discussion of afta data from west greenland 5.6 integration of palaeothermal data with sla 5.7 correlation with the offshore record 5.7.1 the nature of the oligocene hiatus offshore southern west greenland 5.8 tectonic evolution of the west greenland epcm 6. contrasting views on the development of epcms in other areas from landscape studies and thermochronology 6.1 introduction 6.2 southern africa 6.2.1 classic landscape studies 6.2.2 low-temperature thermochronology studies 6.2.3 attempts to reconcile landscape studies and thermochronology in southern africa 6.2.4 possible alternative explanations 6.2.5 evidence for post-break-up subsidence and burial of the southern africa margin 6.2.6 summary 6.3 south-east australia 6.3.1 thermochronology 6.3.2 landscape analysis 6.3.3 attempts to reconcile thermochronology and landscape analysis 6.3.4 evidence for episodic burial and uplift on the south-east australian margin. 6.4 brazil 6.4.1 introduction 6.4.2 studies focused on landscape evidence 6.4.3 studies focused on thermochronology 6.4.4 integration of approaches 6.4.5 continuing controversy 6.4.6 summary 6.5 other areas 6.6 numerical modelling of epcm development 6.7 concluding remarks 7. what processes drive the formation of epcms? 7.1 hypotheses that apply to one or only a few epcms 7.2 permanently uplifted rift margins? 7.3 uplift by underlying hot and/or upwelling mantle 7.4 uplift due to compression 7.5 summary 8. key issues concerning the development of epcms 8.1 steady state or transient landscapes, ancient landscapes or young landscapes? 8.2 landscape evolution and thermochronology: progressive emergence (continual cooling) vs. episodic burial and exhumation (heating and cooling) 8.3. epcms: permanent highs or the result of late uplifts? 8.4 the meaning of “escarpment” 8.5 down-wearing, back-wearing and scale 8.6 are epigene surfaces sometimes re-exposed? 8.7 episodic development of epcms 9. summary and conclusions acknowledgements references geological survey of denmark and greenland bulletin 4, 2003, pp 93-96 93 between october 2001 and the end of 2003 there was a close co-operation between the geological survey of denmark and greenland (geus) and the geological survey department of ghana (gsd), as part of a project to enhance gsd’s institutional capabilities and effectiveness, mainly in the fields of management, geological mapping, map production and data handling. during this period a team of geologists, gis (geographic information system) and database experts as well as administrative staff from geus have visited gsd, and gsd officers have visited geus in copenhagen. the main obstacles to gsd becoming an effective organisation are its status as a department under the ghana ministry of mines, insufficient funding by the government, and poor remuneration of its professional staff. to overcome these obstacles, attempts are being made to change the status of gsd from a ‘civil servant organisation’ into a semiautonomous institution, which will permit the survey to generate funding for its core activities by providing services to outside organisations, and pay better salaries to its personnel. despite many problems, geological mapping has been resumed and three new geological maps have been produced by gsd during the project and stored in gis format. a mapping manual has been prepared, and the structure and ‘mission and vision statements’ for the survey have been revised. geology and mineral resources ghana lies in west africa, at between 5° and 11° north of the equator (fig. 1). it has a surface area of c. 240 000 km2 and c. 20 million inhabitants. its main export commodities are gold, cocoa and timber. apart from gold, ghana hosts major occurrences of manganese, bauxite, diamonds and other raw materials. most of western and northern ghana is underlain by palaeoproterozoic, isoclinally folded, metabasaltic and metasedimentary rocks of the ‘birimian supergroup’ (age ~ 2.2 ga; taylor et al. 1992; hirdes & davis 1998). the metabasalts (fig. 2) form a number of ne–sw-trending volcanic belts, separated by metasedimentary basins (fig. 1; kesse 1985; leube et al. 1990). a slightly younger, unconformable unit of sandstones and conglomerates, the ‘tark-waian group’, overlies the birimian volcanic rocks (fig. 1). rocks of the birimian supergroup are cut by numerous granitoid intrusions (fig. 3), which have yielded ages of 2.2–2.1 ga (hirdes et al. 1992). most of ghana’s mineral deposits are associated with birimian supracrustal rocks; gold also occurs within tarkwaian conglomerates. the central part of the country is occupied by a thick succession of undeformed neoproterozoic to early palaeozoic sedimentary rocks, mainly sandstones (the ‘voltaian basin’), geological survey of denmark and greenland bulletin 4, 93–96 (2004) © geus, 2004 co-operation with the geological survey department of ghana feiko kalsbeek, bjørn hermansen, christian knudsen, leif thorning and marianne thorsen fig. 1. simplified geological map of ghana, modified from kesse (1985). granitoid rocks are subdivided into c. 2.2 ga ‘belt-type granites’ and c. 2.1 ga ‘basin-type granites’, which are mineralogically and chemically distinct (hirdes et al. 1992). which have not been studied in detail. the eastern part of the country consists of a number of thrust sheets composed of sedimentary and crystalline rocks, the ages of which are not well known. they were emplaced during the pan-african orogenic cycle ~ 600 ma ago (hirdes & davis 2002). along the coast and offshore a succession of palaeozoic and mesozoic sedimentary rocks occurs, and there is a limited offshore production of hydrocarbons. the geological survey department of ghana and geus’ involvement the geological survey department of ghana was established in 1913 as the ‘gold coast geological survey’. through the years it has made major contributions to the understanding of the geology of the country, forming the basis for exploration and exploitation of the mineral resources. geological mapping before the second world war was restricted mainly to regional studies, and the first geological map at 1:1 000 000 was published in 1955. after gaining independence in 1957, the new ghana government invested much energy in the survey in order to boost the mineral industry, and during the 1960s and 1970s large parts of the country were geologically mapped in more detail. in the beginning the survey was supported by a number of geologists from the former soviet union, but this co-operation was brought to an end for political reasons. since the early 1980s gsd has been supported in several projects by the german bundesanstalt für geowissenschaften und rohstoffe (bgr). nevertheless, mainly as the result of economic starvation, the activities of the survey during the last few decades have gradually diminished. equipment could not be maintained, qualified staff left, and ultimately the survey was no longer able to perform many of its natural duties. during the late 1990s the ghana government decided that gsd should be gradually reconstructed and upgraded in order to improve its effectiveness. this process has been supported by the world bank, through a loan from the nordic development fund. geus was selected to help with the reconstruction of gsd under a contract with the minerals commission of ghana (ndf credit 156-14). the project ran from 1 october 2001 to 31 december 2003, and the contract sum was c. € 350 000. aims of the project the main topics to be addressed during the project were: 1. formulation of new ‘mission and vision statements’. 2. development of a new organisational structure. 3. reorganisation of the survey’s division for geological mapping, including the preparation of a gsd mapping manual. 4. upgrading the survey’s capability to handle digital data and provide information in digital form. 5. supporting the survey’s programme of human resources development. several of these topics had been initiated by gsd before the beginning of its co-operation with geus. 94 fig. 2. pillow lava at butre, a coastal village west of takoradi. hammer for scale. approach the original plan for the project was that a consultant should provide the survey with a single adviser. in geus’ bid for the project, however, it was suggested that, instead of providing one adviser, a team should be formed to support the survey, consisting of a team leader, who would spend about half of the time in ghana, and a number of short-time specialists would make periodic visits. in this way it was thought that geus would be able to provide support in a number of subjects which could not be covered by one single person. permission was later obtained for gsd officers to visit geus in copenhagen for shorter periods, instead of geus officers visiting ghana. an inception report was prepared jointly by geus and gsd in december 2001 to outline further plans for the project. progress reports were prepared for each of the four following half-year periods, including mission reports on each visit made. in december 2003 a draft final report was submitted. gsd’s formal status gsd is at present a department under the ghana ministry of mines. this has a number of consequences that impede the development of the survey as an effective organisation: 1. governmental funding is insufficient. 2. salaries are very low, as the staff are employed as public servants. young geologists have difficulties in supporting their families and will leave for better paid jobs as soon as possible. as a result of low salaries, poor funding and lack of other incentives, the level of job satisfaction is low. 3. as a ministerial department the survey has to follow strict rules for promotion of its personnel. seniority plays a far greater role in this process than qualifications and performance. 4. dismissal of redundant personnel is virtually impossible under the present situation. the survey has some 300 employees, far more than can be kept usefully occupied with the available funding. the problems for gsd imposed by its status as a ministerial department have been recognised by the government and, with legal advice from geus, a new ‘geological survey act’ has been prepared and submitted to the ministry of mines for further consideration. this act, if and when accepted by the ghana parliament, will transform the survey into a semiautonomous organisation, which will be able to generate funding for its core activities by providing services to outside organisations. moreover, the survey will have the possibility of adjusting the number of staff employed to its actual needs. in a close co-operation between gsd and geus, ‘mission and vision statements’ for this new organisation were formulated, and an organisational structure suggested. main achievements during the project lack of sufficient funding for field work etc. has severely limited the success of the project. nevertheless, with geus support, gsd has made significant progress, especially in the fields of geological mapping, gis applications, database issues and human resources development. geological mapping at the beginning of the present project, geological mapping had been started in five areas in ghana. four of these map95 fig. 3. strongly deformed granitoid gneisses at abuesi, a coastal village west of takoradi. the gneiss has been dated at c. 2180 ma (zircon upb); a granitic vein has yielded zircon and monazite u-pb ages of c. 2100 ma (d.w. davis in loh & hirdes 1999). 96 ping projects were completed by june 2002, and new geological maps were produced. mapping in one area could not be concluded; this is an area in eastern ghana consisting of supracrustal rocks that require experience in stratigraphy and sedimentology, at that time not available at gsd. all mapping teams were visited in the field. most of the geologists responsible for the mapping do an excellent job. however, most of the junior (assistant) geologists need more training before they will be able to carry out independent mapping projects. compilation of the new geological maps (1:100 000) was carried out in accra. thin sections and chemical analyses prepared by geus were used for the correct identification of the various rock types. the final maps were prepared by the staff of gsd’s gis laboratory with support from geus. this was the first time that new gis geological maps have been produced at the survey. digitising old geological maps and reports in order to save old geological maps in gsd’s archives from deterioration, digitising of existing maps has had a high priority during the project. a database was made with information on all existing maps. geus helped train gsd staff in the application of gis in map production. fourteen old maps were digitised during the project period, and a cdrom with these maps has been prepared. this has been one of the most successful parts of the project. databases since there is an urgent need for gsd to be able to present its data to the public in the form of databases, a major effort has been made to upgrade the survey’s experience in this field. geus has analysed which databases would be most needed and, in order to improve the understanding of database issues at gsd, a minerals occurrence database has been established. human resources development continuous training of its staff as well as recruitment of skilled young geologists will be necessary for gsd to become a strong and professional geological survey. in this context major parts of geus’ efforts under the project have been directed to practical on-the-job training of gsd personnel. apart from training related to geological mapping, gis and databases described above, geus has provided a collection of thin sections with descriptions of typical greenland rocks, given courses in the microscopic study and classification of rocks, provided training in geological mapping for two gsd geologists as participants in geus’ 2002 mapping project in greenland, sponsored an excursion for a group of gsd geologists in the western and central regions of ghana, and conducted a training course for gsd’s management on internal co-operation and related issues. sustainability there is no doubt that co-operation with geus has improved the ability of gsd to carry out several of its main functions. however, for these improvements to be sustainable in the future, it is imperative that the survey is changed into a semi-autonomous organisation. only then will the survey have the possibility to generate income to cover parts of its running expenses. moreover, it would be able to adapt the number of staff to its actual needs and provide better wages for its remaining personnel, thereby improving job satisfaction and limiting the number of well-qualified staff that leave the organisation. references hirdes, w. & davis, d.w. 1998: first u-pb zircon age of extrusive volcanism in the birimian supergroup of ghana/west africa. journal of african earth sciences 27, 291–294. hirdes, w. & davis, d.w. 2002: u-pb zircon and rutile metamorphic ages of dahomeyan garnet-hornblende gneisses in southeastern ghana, west africa. journal of african earth sciences 35, 445–449. hirdes, w., davis, d.w. & eisenlohr, b.n. 1992: reassessment of proterozoic granitoid ages in ghana on the basis of u/pb zircon and monazite dating. precambrian research 56, 89–96. kesse, g.o. 1985: the mineral and rock resources of ghana. xiv + 610 pp. rotterdam: balkema. leube, a., hirdes, w., mauer, r. & kesse, g.o. 1990: the early proterozoic birimian supergroup of ghana and some aspects of its associated gold mineralization. precambrian research 46, 139–165. loh, g. & hirdes, w. 1999: explanatory notes for the geological map of southwest ghana 1:100,000 sekondi (0402a) and axim (0403b) sheets. with contributions from c. anani, d.w. davis and u.k. vetter. geologisches jahrbuch reihe b. heft 93, 149 pp. + 2 maps. taylor, p.n., moorbath, s., leube, a. & hirdes, w. 1992: early proterozoic crustal evolution in the birimian of ghana: constraints from geochronology and isotope geology. precambrian research 56, 97–111. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ckn@geus.dk << /ascii85encodepages false /allowtransparency false /autopositionepsfiles true /autorotatepages /all /binding /left /calgrayprofile (dot gain 20%) /calrgbprofile (srgb iec61966-2.1) /calcmykprofile (u.s. web coated \050swop\051 v2) /srgbprofile (srgb iec61966-2.1) /cannotembedfontpolicy /warning /compatibilitylevel 1.4 /compressobjects /tags /compresspages true /convertimagestoindexed true /passthroughjpegimages true /createjdffile false /createjobticket false /defaultrenderingintent /default /detectblends true 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<feff4f7f752890194e9b8a2d5b9a5efa7acb76840020005000440046002065874ef65305542b8f039ad876845f7150cf89e367905ea6ff0c4fbf65bc63d066075217537054c18cea3002005000440046002065874ef653ef4ee54f7f75280020004100630072006f0062006100740020548c002000520065006100640065007200200035002e0030002053ca66f465b07248672c4f86958b555f3002> >> >> setdistillerparams << /hwresolution [2400 2400] /pagesize [595.000 842.000] >> setpagedevice geological survey of denmark and greenland bulletin 13, 2007, 69-72 las tablas de daimiel, together with other wetlands in la mancha, spain, situated in the upper guadiana basin (fig. 1), has been catalogued as a biosphere reserve area since 1981 as part of the unesco man and the biosphere programme. between the mid-1970s and late 1980s, over 150 000 hectares of new irrigation areas were established, mainly as a result of private initiative. the average recharge rate of groundwater in the western la mancha aquifer in the upper guadiana basin is estimated to be between 200 and 500 million m3 per year, in dry and wet years respectively. re charge also depends on the depth of the water table (martínez-cortina & cruces 2003). abstraction reached 600 million m3 per year by the end of the 1980s. up to this time a total of 3000–5000 million m3 of the upper gua diana basin aquifer’s water reserves was withdrawn (bromley et al. 2000; lopéz-geta et al. 2006). the intensive use of groundwater has been a main factor for the improvement of the social and economic situation in this region, with a population of about half a million people, and where the agricultural sector is very important (llamas et al. 2006). water-table drawdown due to the intensive abstraction of groundwater for irrigation has caused severe negative impacts on wetlands, streams and rivers, and has resulted in a lowering of groundwater levels by up to 50 m. the main conflicts in the area are between farmers and conservationists, between central, regional and local government water agencies, and between small farmers and big farmers. the conflicts began about three decades ago (llamas 1988) and have not yet been settled. in 2001 the spanish parliament asked the govern ment to present a hydrological plan for the upper guadiana basin within one year. more than 20 draft proposals have been presented, the last one in 2006 with a budget of almost four billion euros. this proposal has been met with strong opposition from most farmer lobbies. the guadiana basin is one of seven transboundary case studies of the eu newater research project (new approaches for adaptive water management under uncertainty). the prin cipal water-management issues in the project are addressed by adaptive and integrated water-resource manage ment. this includes uncertainty and risk mitigation, gov er n ance, crosssectoral integration, scale analysis, information management, stakeholder participation, financial aspects, system resilience and vulnerability. one work block in the newater project has the task of translating research outputs into tools for practitioners and end-users. as part of this effort, bayesian belief networks (bns) were selected as one possible tool to be developed as an aid to stakeholder participation in integrated assessment of gaps, being a suitable tool for dialogue in order to identify gaps in water-resource management functions, gaps to meet the goals of the eu water framework directive and to analyse management potentials and constraints. the purpose of this paper is to describe the testing of bns as a tool for participatory integrated assessment and adaptive and integrated water-resource management in the upper guadiana basin. participatory integrated assessment participatory integrated assessment can be considered a form of participatory policy analysis, which aims to support the policy process by designing and facilitating policy debate and argument. assessment is integrated when it draws on a broader set of knowledge domains than are represented in the bayesian belief networks as a tool for participatory integrated assessment and adaptive groundwater management: the upper guadiana basin, spain hans jørgen henriksen, per rasmussen, john bromley, africa de la hera portillo and m. ramón llamas © geus, 2007. geological survey of denmark and greenland bulletin 13, 69–72. available at: www.geus.dk/publications/bull 69 fig. 1. location of the guadiana basin in spain. the upper guadiana basin includes las tablas daimiel and upstream areas in the castilla-la mancha region. major rivers indicated. research product of a single discipline. assessment is distinguished from disciplinary research by its purpose: to inform policy and decision-making, rather than to advance knowledge for its intrinsic value (hisschemöller et al. 2001). a wide range of methods and techniques can be drawn from social psychology, policy sciences, decision analysis and anthropology (hisschemöller et al. 2001) for high-level participatory integrated assessment. some of these, like brainstorming or decision seminars, although well established, are of limited value for integrated water-resource management because a proper understanding of the spatial and temporal variation and the complexity within river basins requires a modelling approach (croke et al. 2007). according to jakeman & letcher (2003) the tools for participatory integrated assessment must: 1. be problem-focussed, using an iterative, adaptive ap proach that links research to policy; 2. possess an interactive, transparent framework that en hances communication; 3. be enriched by stakeholder involvement and dedicated to adoption; 4. connect complexities between the natural and human en vironment, recognising spatial dependencies, feedbacks and impediments; and 5. attempt to recognise essential lacking knowledge. jakeman & letcher (2003) list several tools for participatory integrated assessment, e.g. system dynamics, bns, metamodels, risk assessment approaches, coupled component models, agent-based models and expert systems. here bns are in focus as a tool for adaptive and integrated water management in the upper guadiana basin test case. bayesian belief networks a bayesian belief network (bn) is a type of decision support system based on probability theory which implements bayes’ rule of probability (jensen 2002; bromley 2005). this rule shows mathematically how existing beliefs can be modified with the input of new evidence. bns organise the body of knowledge in a given area by mapping out relationships among key variables and encoding them with numbers that represent the extent to which one variable is likely to affect another. bns have gained a reputation for being a powerful technique to model complex problems involving uncertain knowl edge and uncertain impacts of causes. ideally, bns are a technique to assist decision-making that is especially helpful when there is scarcity and uncertainty in the data used in making the decision and the factors are highly interlinked, all of which makes the problem very complex. the graphical nature of bns facilitates formal discussion of the structure of the proposed model. furthermore, the ability of bns to describe the uncertain relationships between variables is ideal to describe the relationship between events, which may not be well understood. bns help water managers, stakeholders and scientists (1) to visualise and recognise, in the face of complexity and uncertainty, the relationships between different actions and consequences; (2) to make learning about water-resource systems more efficient; and (3) to encourage the involvement of social and political values in water-resource management (e.g. henriksen et al. 2007a, b). furthermore, it has been judged that bns are an excellent tool for integrating different domains, e.g. socio-economy, hydrology and groundwater quality data of different knowledge types (monitoring data, models and expert opinions; henriksen et al. 2007a). here the guidelines from the merit project (bromley 2005) can help support a successful and efficient involvement of stakeholders in the participatory integrated assessment process, a process which is demanding to run due to multiple frames and opposing interests. design for testing the enhanced bayesian belief network tool a test of an enhanced bn tool is being undertaken in the upper guadiana basin as part of the newater project. the test involves the construction of a bn to represent the management of groundwater levels in the region, taking into account the social, economic, hydrological and ecological consequences of alternative irrigation and groundwater management scenarios (table 1). in november 2006, an initial workshop was held at the geological survey of spain (igme) in madrid with participants from the case study group. during this workshop a preliminary bn for the upper guadiana basin was developed by bn experts from igme, the university complutense de madrid, the geological survey of denmark and greenland, and the centre of ecology and hydrology, wallingford, uk, 70 together with a representative of the water managers of the basin responsible for water planning in relation to implementation of the eu water framework directive. a joint workshop with all stakeholders to finalise the network has been planned for the first half of 2007. the process and method for constructing the bn in the upper guadiana basin test will follow the merit guidelines (bromley 2005). in the following we present the preliminary bn and the hypotheses relating to the use of the tool in participatory integrated assessment and adaptive management in the guadiana basin. results of testing bayesian belief networks for adaptive water management the initial step in network design was to establish the space and time boundaries of the system being modelled. it was agreed to restrict the model to the upper guadiana basin, and a one-year time period for groundwater level and socioeconomic consequences was decided. the pilot bn for the upper guadiana basin case which emerged from this process is shown in fig. 2. the network deals with the way in which different management actions influence irrigation water use, groundwater level, crop pattern, farmers’ income, wetland recovery, productivity and employment in the region (fig. 2). included among the potential actions that might be taken are: (1) acquisition of water rights; (2) law enforcement; (3) common agricultural programmes (cap) subsidies; and (4) annual management plans. climate and the initial state of the aquifer are included as control factors. the indicators (objectives) in the network include: (1) groundwater levels; (2) impact on wetland recovery; (3) agricultural productivity; (4) farmers’ income; and (5) levels of employment in the region. when running the bn, combinations of actions can be selected and calculated. it is hypothesised that bns fully support four of the five requirements proposed by jakeman & letcher (2003) for participatory integrated assessment (table 2). one requirement, the representation of spatial dependencies, is only partly supported (e.g. input to the decision-making about which specific wetlands that will be recovered by a certain increase in groundwater level has to be evaluated using a groundwater model). however, as stated by pascual (2005): “the beauty of bns lies in their explanatory power: observations about any node generates knowledge about all other nodes, providing one with a tool to draw transparent, ration al inferences in a probabilistic world”. this illustrates that the tool can be used for diagnosis and social learning. bns allow targeted modelling, participatory integrated assessment and strong support for sense and decision-making in cases with multiple frames (e.g. when stakeholders perceive their environment differently, and frame and construct their world in different ways) that create ambiguous situations and conflicting interests hindering sustainable solutions for man71 fig. 2. preliminary bayesian network for the upper guadiana basin. the objectives of the bayesian network are to analyse the way in which different management actions will influence irrigation water use, change in groundwater level, crop pattern, farmers’ income, wetland recovery, productivity and employment. agement of the environment. the tool and the probability tables (numbers) are not easily understood if not properly explained. thus, training and introduction to the tool and the statistical background behind bns is important (table 2). acknowledgements the work reported from the newater project has been financially supported by the european commission under contract number 511179 (goce). integrated project in priority 6.3 global change and ecosystems in the 6th eu framework programme. references bromley, j. 2005: guidelines for the use of bayesian networks as a participatory tool for water resource management, 117 pp. wallingford: centre for ecology and hydrology. bromley, j., cruces, j., acreman, m., martinez, l. & llamas, m.r. 2000: groundwater over-exploitation in the upper guadiana catchment, central spain: the problems of sustainable groundwater resources man agement. international water resources development 17, 379–396. croke, b.f.w., ticehurst, j.l., letcher, r.a., norton, j.p., newham, t.t.h. & jakeman, a.j. 2007: integrated assessment of water resources: australian experiences. water resource management 21, 351–373. henriksen, h.j., rasmussen, p., brandt, g., bülow, d.v. & jensen, f.v. 2007a: engaging stakeholders in construction and validation of baye sian belief network for groundwater protection. in: castelletti, a.e.r. & soncini-sessa, r. (eds): topics on system analysis and integrated water resource management, 49–72. amsterdam: elsevier. henriksen, h.j., rasmussen, p., brandt, g., bülow, d.v. & jensen, f.v. 2007b: public participation modelling using bayesian networks in management of groundwater contamination. environmental modell ing & software 22, 1101–1113. hisschemöller, m., tol, r.s.j. & vellinga, p. 2001: the relevance of participatory approaches in integrated environmental assessment. inte grated assessment 2, 57–72. jakeman, a.j. & letcher, r.a. 2003: integrated assessment and modelling: features, principles and examples for catchment management. envi ronmental modelling and software 18, 491–501. jensen, f. 2002: bayesian networks and decision graphs: statistics for engineering and information science, 296 pp. new york: springerverlag. llamas, m.r. 1988: conflicts between wetland conservation and groundwater exploitation: two case histories in spain. environmental geol ogy and water sciences 11, 241–251. llamas, m.r., martínez-santos, p. & hera, a. de la 2006: dimensions of sustainability in regard to groundwater resources: an overview. pro ceedings of the international symposium on groundwater sustain ability, alicante, spain, 24–27 january 2006, 1–13. madrid: instituto geológico y minero de españa. lópez-geta, j.a., fornés, j.m., ramos, g. & villarroya, f. 2006: groundwater. a natural underground resource, 107 pp. madrid: insti tuto geológico y minero de españa, unesco and fundación mar celino botín. martínez-cortina, l. & cruces, j. 2003: the analysis of the intensive use of groundwater in the upper guadiana basin (spain) using a numerical model. in: llamas, m.r. & custodio, e. (eds): intensive use of groundwater, challenges and opportunities, 285–294. london: taylor and francis. pascual, p. 2005: wresting environmental decisions from an uncertain world. environmental law institute news and analysis 8-2005, 10539–10549. authors’ addresses h.j.h. & p.r., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: hjh@geus.dk j.b., oxford university centre for the environment, south parks road, oxford ox1 3qy, uk. a.d.l.h.p., geological survey of spain, 23 rios rosas, itge-e 28003 madrid, spain. m.r.l., university of complutense, ciudad universitaria, 28040 madrid, spain. 72 geological survey of denmark and greenland bulletin 35, 2016, 87-90 87© 2016 geus. geological survey of denmark and greenland bulletin 35, 87–90. open access: www.geus.dk/publications/bull the concept of utilising available pore space in deep saline sandstone aquifers for storage of co2 was recognised in the late 1980s. in 1996, the first commercial co2 storage project began with injection into sandstones of the utsira formation in norway. the formation is located above the sleipner formation from where the sleipner field produces natural gas. the project was initiated due to a high co2 content of the natural gas, which was subjected to a norwegian offshore carbon tax. the natural gas is produced on the sleipner platform where the co2 is separated, captured and reinjected from a neighbouring platform. the potential for using the technology to reduce co2 emissions from large stationary point sources initiated many research projects aimed at mapping areas with potential co2 storage capacity around the world. in 2008, the nordic countries decided to set up a special venture for climate, energy and the environment by launching the top-level research initiative promoting research within six sub-programmes, including one on carbon, capture and storage (ccs). with this background the nordic ccs competence centre (nordiccs) was initiated in 2011, involving major nordic ccs research institutes, industry and stakeholders. one of the main outcomes of the project, which terminated in 2015, is a webbased nordic co2 storage atlas (data.geus.dk/nordiccs/ map.xhtml), which aims to make ccs-related data and interpretations available to decision makers and the public. the newly released atlas combines data from previous co2 storage screening and mapping projects (gestco, eu geocapacity and the norwegian co2 storage atlas) with new data for areas not previously covered. the atlas gives an overview of storage options and the associated reservoir properties in denmark, norway, sweden and iceland. co2 storage site screening in the nordic region the project focused on co2 storage (1) in sandstone aquifers, (2) by chemical reaction in basalts and (3) in depleted hydrocarbon fields. large-scale geological storage of co2 in sandstone aquifers requires the presence of a porous and permeable subsurface layer, a burial depth of minimum 800 m to keep the co2 as a dense phase, and an adequate top seal preventing the buoyant co2 from migrating to the surface. areas with the largest storage potential are associated with sedimentary basins containing widespread sandstone layers. sedimentary basins with storage potential are situated as a marginal belt around the scandinavian peninsula from the baltic sea, through denmark and along the norwegian coast, whereas the shallow sedimentary basins in finland are not considered appropriate for co2 storage (teir et al. 2010). in iceland, the storage potential is not related to sedimentary basins, but to injection of co2saturated water into porous basalts (fig. 1). the dissolved co2 reacts with divalent cations in the basalt-forming stable carbonate minerals such as calcite, dolomite, magnesite, siderite, and mg-fe carbonate solid solutions (gislason et al. 2010). compared to storage in saline aquifers, the co2 storage capacity in hydrocarbon fields is in general minor, but latemapping of the co2 storage potential in the nordic region karen lyng anthonsen, peter frykman and carsten møller nielsen 500 km n wells (felicia-1, j-1, k-1) hanstholm structure sediments < 540 ma caledonian baltic shield (fennoscandia) primarily basalts iceland k-1 j-1felicia-1 norway finland sweden denmark fig. 1. generalised geological overview of the precambrian baltic shield and the sedimentary basins. 8888 stage oil production may profit from enhanced oil recovery (eor) by injection of co2. thus the highest potential for industrial-scale storage is obtained when eor is used in connection with the co2 storage. the geological and engineering knowledge accumulated from producing oil and gas fields is also more detailed than for aquifer storage sites. the data compiled in the web-based atlas include an approximate outline of storage formations and aquifers with information for each unit on depth, thickness, lithology, proportion of sand, age and reservoir type, based on well data and seismic interpretations. furthermore, reservoir properties such as porosity, permeability, salinity, co2 density, storage efficiency factor and estimated storage capacity are listed. in order to illustrate the geological complexity, major faults are included in the atlas. likewise, an outline of the sealing formations is included in order to indicate storage integrity. the compiled data were used to characterise and rank storage formations and traps and to calculate co2 storage capacities. the ranking criteria were grouped into four main categories: reservoir properties, seal properties, safety/ risk and maturity/data coverage. this resulted in a selection of the most prospective nordic storage areas, based on available geological knowledge up to 2014 (anthonsen et al. 2014). the ranking revealed that the most prospective storage areas are found in the norwegian north sea. this is basically a result of the knowledge from the intensive oil and gas exploration making the norwegian areas more mature for exploitation of storage capacity. it has to be stressed that there are large uncertainties in many of the evaluated parameters and that more data and further data analysis are required before any of the sites are ready for co2 injection. mapped co2 storage capacity the storage capacity estimates make use of the same methodology as the eu geocapacity project; see vangkildepedersen et al. (2009). the total mapped co2 storage capacity for denmark, norway and sweden is 134 000 megaton (mt). the storage capacity related to saline aquifers is 120 000 mt, with 22 000 mt in denmark, 94 600 mt in norway (hereof 72 800 mt in the north sea) and 3 400 mt in sweden. the total number includes 14 000 mt in hydrocarbon fields, with 2 000 mt in denmark and 12 000 mt in norway (røkke et al. in press; fig. 2). it should be emphasised that the presented storage capacities are regarded as simple estimates based on volumetric calculations of the available pore space and multiplied with a storage efficiency factor. improved geological data and reservoir modelling work will be necessary to narrow the specific uncertainties for the storage capacity estimates. storage capacity estimates for porous basalts are based on a different methodology, and for onshore iceland the calculated capacity ranges between 21 000 and 400 000 mt depending on the calculation approach (snæbjörnsdóttir et al. 2014). the storage capacity has to be seen in relation to the co2 emission, which was 152.8 mt in 2011 for all large stationary point sources (emission >70 kt) in the five nordic countries, with 25.8 mt in denmark, 55.6 mt in finland, 1.6 mt in iceland, 23.1 mt in norway and 46.6 mt in sweden. on a european scale the total emission from stationary point sources mapped in the eu geocapacity project was 2 000 mt (anthonsen et al. 2011), implying that theoretically all co2 emissions for 70 years from these sources in europe could be stored in the nordic region. modelling co2 storage capacity in denmark the procedure for the estimation of storage capacity is illustrated by a case study of the hanstholm structure, using simple, static calculations supplemented with dynamic simulations. the dynamic simulation has the advantage that both reservoir properties such as heterogeneity and maritime borders sweden finland norway iceland denmark hydrocarbon fields selected prospective storage areas mapped formations with storage potential porous basalts < 0.8 mill. years sedimentary basins 500 km n fig. 2. all mapped nordic co2 storage formations (blue) and the selected most prospective areas in green. the dark blue area in iceland is the highly porous basalt areas considered most promising for co2 injection. 89 operational conditions can be accounted for, leading to a more realistic capacity number. the dynamic assessment requires that a 3d reservoir model is constructed, as it involves flow and pressure calculations over the time span of the operational period. the background for constructing this model is briefly described here. the informal name hanstholm structure is used for an offshore domal closure covering 603 km2, situated offshore c. 40 km north-west of the city of hanstholm (fig. 1). the water depth at the site is c. 30 m. the target for storage is upper triassic – lower jurassic sandstones of the gassum formation. this formation consists of fineto mediumgrained, locally coarse-grained sandstones interbedded with heteroliths, claystones and locally thin coal beds (michelsen et al. 2003; nielsen 2003). the sandstones were deposited by repeated progradation of shoreface and deltaic units forming laterally continuous sheet sandstones separated by offshore marine claystones. fluvial sandstones dominate in the lower part of the formation as in most of the fennoscandian border zone. the structure is situated close to the edge of the fjerritslev fault of the sorgenfrei-tornquist zone, and is formed by uplift due to post-depositional salt tectonics. the structure is interpreted from the depth structure map of the ‘top triassic’ as defined by japsen & langtofte (1991), and has been used as a template for defining top and bottom of a reservoir with uniform thickness. the depth to top reservoir is approximately 890 m below mean sea level, and the deepest closing contour is at approximately 1330 m (fig. 3a). the theoretical spill point is situated at the south-eastern flank of the structure spilling into the thisted domal structure. the structure has not been drilled and data for the reservoir have to be extrapolated from information from the nearby felicia-1, j-1 and k-1 wells (fig. 1). it should be noted, however, that felicia-1 is drilled at the crest of a rotated fault block, and is believed to show an extraordinarily large thickness of the gassum formation with a thick mudstone in the middle part, reflecting topographic influence from the nearby salt pillow during deposition. this may result in marked differences in reservoir properties between this well and the undrilled structure. the well j-1 some 30–40 km to the north-east has therefore been used as a template for the sand-shale sequence in the reservoir model. the claystones of the fjerritslev formation form the top seal of the aquifer. the fjerritslev formation is expected to be c. 500 m thick above the hanstholm aquifer. the reservoir model was used for a capacity study by simulating seven injection wells around the perimeter of the flank, and their positions were optimised by iteration to give the most complete filling pattern on the structure (fig. 3b). co2 was injected at a constant rate of 4.2 mt per well per year. the dynamic simulations account for an injection period of 40 years and with a preconditioned safety 4 km4 km n-1000 top reservoir elevation depth [m] -1400 -1800 -2200 0.0 0.1 0.2 0.3 0.4 0.5 sg n a b fig. 3. a: maps with depth contours for the top reservoir level and the position of the seven injection wells. the seven wells were positioned by iteration to effectively fill most of the structure. b: an optimum filling simulation of the hanstholm structure. injection period is 40 years. 0 1000 2000 3000 4000 5000 6000 7000 dynamicstatic 40% efficiencystatic 100% efficiency m to n c ap ac it y hanstholm capacity estimates fig. 4. comparison of different capacity estimates including values from the dynamic simulation of optimum filling of the hanstholm structure in denmark. the previous estimate was based on static calculations and an assumed effieciency factor of 40%. 9090 margin of the allowed pressure increase set to maximum 85% of the lithostatic pressure below the cap rock in order to avoid any fracture propagation in the cap rock. the resultant storage capacity was 1170 mt co2, but a change of the safety margin will naturally influence the storage capacity. the derived dynamic capacity for the hanstholm structure can be compared to previous estimates of 2753 mt based on the static model characteristics and an assumed efficiency factor; in this case 40% efficiency (fig. 4) as described by larsen et al. (2003). summary mapping of sandstone units (aquifers) and their associated reservoir properties have resulted in a web-based co2 storage atlas. the reservoir data and properties were used to characterise and rank the potential storage areas and sites in an attempt to point out the most prospective ones based on currently available geological knowledge (fig. 2). it is essential for the development of a co2 storage site to know how much capacity is available. in the initial screening phase static theoretical estimates are used, but dynamic modelling of co2 injection is very important in order to narrow the uncertainties of the storage capacities. the reduction in total storage capacity from previously published static calculations to the modelled dynamic calculations is one of the key conclusions from the co2 injection simulation. however, even if a reduction of the static capacity estimates is taken into account, it is clear that the nordic region has substantial storage capacity in saline aquifers (anthonsen et al. 2014; lothe et al. 2015). acknowledgements this article has been produced with support from nordiccs, under the top-level research initiative co2 capture and storage programme, and nordic innovation. the authors acknowledge the following partners for their contributions: statoil, gassco, norcem, reykjavik energy, co2 technology centre mongstad, vattenfall and the top-level research initiative (project number 11029). references anthonsen, k.l., frykman, p. & nielsen, l.h. 2011: the potential for geological storage of co2 in denmark is very promising. in: sønderberg petersen, l. & larsen h. (eds): energy systems and technologies for the coming century. risø international energy conference 2011, may 10–12. proceedings, 48–55. anthonsen, k.l., aagaard, p., bergmo, p.e.s., gislason, s.r., lothe, a.e., mortensen, g.m. & snæbjörnsdottir, s.ó. 2014: characterisation and selection of the most prospective co2 storage sites in the nordic region. energy procedia 63, 4884–4896. gislason, s.r., wolff-boenisch, d., stefansson, a., oelkers, e.h., gunnlaugsson, e., sigurdardottir, h. & sigfussson, b. 2010: mineral sequestration of carbon dioxide in basalt: a pre-injection overview of the carbfix project. international journal of greenhouse gas control 4(3), 537–545. japsen, p. & langtofte, c. 1991: geological map of denmark 1:400 000. the danish basin: ‘top trias’ and the jurassic – lower cretaceous. danmarks geologiske undersøgelse map series 30. larsen m., bidstrup t. & dalhoff, f. 2003: mapping of deep saline aquifers in denmark with potential for future co2 storage. a gestco contribution. danmarks og grønlands geologiske undersøgelse rapport 2003/39, 83 pp. lothe, a.e., emmel, b., bergmo, p.e., mortensen, g.m. & frykman, p. 2015: updated estimate of storage capacity and evaluation of seal for selected aquifers (d26). nordiccs technical report d 6.3.1401, 80 pp. michelsen, o., nielsen, l.h., johannessen, p.n., andsbjerg, j. & surlyk, f. 2003: jurassic lithostratigraphy and stratigraphic development onshore and offshore denmark. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 147–216. nielsen, l.h. 2003: late triassic – jurassic development of the danish basin and the fennoscandian border zone, southern scandinavia. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 459–526. røkke, n.a., aarlien, r., mazzetti, m., kielland haug, j.j., skagestad, r., onaeheim, k., lund, h., kjärstad, j. & anthonsen, k.l. in press: final report. nordiccs, nordic ccs competence centre. nordic innovation publication. snæbjörnsdóttir, s.ó., wiese, f., fridriksson, t., ármannsson, h., einarsson, g.m. & gislason, s.r. 2014: co2 storage potential of basaltic rocks in iceland and the oceanic ridges. energy procedia 63, 4585–4600. teir, s. et al. 2010: potential for carbon capture and storage (ccs) in the nordic region. vtt tiedotteita – research notes 2556, 53–73. vangkilde-pedersen, t. et al. 2009: assessing european capacity for geological storage of carbon dioxide – the eu geocapacity project. energy procedia 1, 2663–2670. authors’ address geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: kla@geus.dk geological survey of denmark and greenland bulletin 20, 2010, 99–102 99 in the c. 40 000 km2 large phu quoc basin south-west of vietnam reflection seismic data suggest a thin-skinned thrust-fault complex concealed by neogene marine sediments (fig. 1; fyhn et al. 2010). the deformed sedimentary succession in the complex is of early cretaceous age, which is documented by biostratigraphical studies of outcrops and a 500 m deep well on the phu quoc island. a model for the thrust-fault deformation suggests that piggy-back basins were formed during displacement along the thrust faults. the translation was 3–8 km from east to west. the model is based on detailed structural analyses of 36 seismic sections that cover the phu quoc basin (fig. 1). the main structural elements in the complex are flats and ramps with hanging-wall anticlines developed above the ramps. the crests of the hanging-wall anticlines occur as remnants of partially eroded structural highs. this paper describes the thin-skinned thrust-fault structures that form the basis for the interpretation of the concealed fold-belt complex in the phu quoc basin. architectural framework of the fold belt along the west coast of vietnam folded and thrust-deformed mesozoic and palaeozoic sedimentary rocks form a n–sstriking fold belt referred to as the kampot fold belt (fyhn et al. 2010). to the north the fold belt continues into a hilly area at the border between vietnam and cambodia, where it forms a mountain range. to the south, offshore the west coast of vietnam, the fold belt is concealed below neogene marine sediments. from west to east the fold belt is divided into a distal, intermediate and proximal part (fig. 1). the distal part is a frontal wedge that passes into the foreland about 100 km west of the vietnam coast. the intermediate part is characterised by moderately folded, hanging-wall anticlines with thrust displacements in the order of a few kilometres. the transition from the intermediate to the proximal parts is located close to the nam du archipelago (pedersen et al. 2009). two interpreted, representative seismic sections have been selected to illustrate the architectural style in structural cross-sections perpendicular to the main trend of the thrustfault belt (figs 2, 3). phu quoc p ro x im al i m b ri ca te p ar t 104˚00'103˚40' 104˚20'e20 km ■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ 9˚20' fig. 3 fig. 2 9˚00' 8˚40' 10˚00'n ▲ ■■ distal part intermediate part proximal part thrust fault footwall syncline flexure on ramp dip conventional seismic lines high resolution seismic lines enreca-2 well hanging-wall anticline 500 km laos thailand cambodia vietnam nam du fig. 1. map of the study area off sw vietnam, showing the seismic grid which formed the basis for the study of the phu quoc basin. on the map the main structural elements are shown with the trends of anticlines and thrust faults interpreted from the seismic sections. the location of the 500 m deep well enreca-2 on the phu quoc island and the position of the two cross-sections in figs 2 and 3 representing the distal and the intermediate parts of the deformation complex are also shown. the inset map shows the location of the investigated area in se asia. © geus, 2010. geological survey of denmark and greenland bulletin 20, 99–102. open access: www.geus.dk/publications/bull thin-skinned thrust-fault tectonics offshore south-west vietnam stig a. schack pedersen, lars ole boldreel, emil bach madsen, mette bjerkvig filtenborg and lars henrik nielsen 100100 the distal part of the tectonic complex the distal part of the complex passes gradually into the undeformed foreland where almost horizontally bedded cretaceous sediments occur, which are separated from overlying neogene marine sediments by a distinct unconformity. thrust-faults dipping <2° are found from the foreland towards the intermediate part. a number of minor ramps are present until the main décollement zone passes into a deeper level along a moderately dipping ramp, above which the westernmost, major hanging-wall anticline is seen (fig. 2). the main décollement zone is located below 2.5 km of sediments in the distal area, west of the foreland-near, hanging-wall anticline. the ramp takes the décollement zone down to 3 km below surface. at this depth the resolution of the seismic data becomes low, so no further interpretations are carried out. the hanging-wall anticline is relatively flat-topped, which corresponds well with a model for hanging-wall anticlines where the displacement is about half the thickness of the thrust-faulted sedimentary rock unit (pedersen 2006). the western limb of the hanging-wall anticline dips c. 25°w, and the uppermost depositional unit below the neogene unconformity on-laps the dipping limb. the beds in this unit are gently tilted towards the west due to the main, gentle dip of the depositional wedge in the frontal part of the thrust system. therefore this uppermost depositional unit may be regarded as a piggy-back basin, which was formed during the translation of the frontal part of the thrust-fault complex. interpretations of the seismic sections north of the presented cross-section show additional examples of piggyback basins. the most important of these is a basin found south-east of the phu quoc island. it formed between two hanging-wall anticlines during their fold and translation development. in addition, we have interpreted the gently deformed lower cretaceous deposits that were penetrated by the enreca-2 well on the phu quoc island to continue into the deformed succession offshore. in our tectonic model the environment surrounding phu quoc island is regarded as located at the transition from the distal to the intermediate part of the tectonic complex. the intermediate part of the tectonic complex two or more levels of thrust faults with flats and ramps have developed in the intermediate part of the complex (figs 1, 3). the number of ramps increases eastwards, which leads to an increasing number of hanging-wall anticlines, some of which are interpreted to have developed into antiformal stacks. the appearance of duplex structures initially lifted the top of the tectonic complex to a higher level. however, this is only recognised as a deeper level of erosion into the tectono-stratigraphic units. the displacement increases to 3–8 km resulting in some variation in the structural style of neogene marine sediments lower décollement surface pre-cretaceous deposits neogene unconformity lower cretaceous terrestrial deposits thrust fault hwr fwr 1 km cretaceous foreland-basin unconformity 0 500 1000 1500 2000 2500 3000 3500 0 500 1000 1500 2000 2500 3000 3500 ww ew e tw o -w ay t ra ve l ti m e (m se c) tw o -w ay t ra ve l ti m e (m se c) c. 1 km fig. 2. seismic section representing the distal part of the tectonic complex. in the western part of the cross-section the most distally appearing hanging-wall anticline is seen, and to the west the almost planar horizontal bedding extends into the foreland basin. beds in the youngest phu quoc unit onlap the moderately dipping beds on the western limb of the hanging-wall anticline, and these beds as well as the top of the hanging-wall anticline are truncated by the neogene unconformity. hwr: hanging-wall ramp, fwr: footwall ramp. blue lines: prominent bedding surfaces. red lines: thrust faults. purple line: the main neogene unconformity that truncates the structures in the tectonic complex. for location see fig. 1. 101 the hanging-wall anticlines. at the transition from the intermediate to the proximal part, erosional remnants of the hanging-wall anticlines are preserved as scattered islands rising a few hundred metres above sea level. the easternmost island in the nam du archipelago provides an example of this. on this island the thrust deformation elevated permian rhyolitic hyaloclastics to a surface-near position (pedersen et al. 2009). the proximal part of the tectonic complex the proximal part of the tectonic complex is only covered by seismic data only in the south-eastern part of the study area (fig. 1). here steeply dipping imbricate structures with 200–500 m thick thrust sheets occur. the depth to the décollement zone is more than 3 km. onshore the imbricate structures in the proximal part are exposed in the mountain range at the border between vietnam and cambodia. in this area the general dip of the thrust sheets is about 30°, and the deformed sedimentary rocks comprise upper palaeozoic sandstones and shales, permian carbonates and triassic sandstones, arkoses and conglomerates. in the thrust-fault zones, shearing and low-grade metamorphism have altered the sediments with recrystallisation of albite and formation of chlorite and biotite, corresponding to lower to medium greenschist facies. no minerals indicating higher metamorphic grades have been recognised. one of the small islands c. 50 km north-north-east of nam du is located in the proximal part. on this island, permian carbonates are thrustfaulted over jurassic shales and sandstones. a granitic plug occurs in the middle part of the island and granitic sills have intruded the jurassic succession. the intrusions are of lower cretaceous age (pedersen et al. 2009). cretaceous granitic intrusions occur on the mainland east of the proximal part of the tectonic complex. they form a cretaceous magmatic arc in the hinterland zone of the kampot fold belt (fyhn et al. 2010). the intrusive age of the granites is based on radiometric dating, and the dating of the uplift and erosion is based on fission-track studies that indicate exhumation during the eocene (fyhn et al. 2010). discussion the timing of the deposition of the basin fill and the deformation of the phu quoc basin are addressed by the tectonic model. the structural interpretation of the thrust-fault structures implies that the piggy-back basins formed in the cretaceous, which in turn suggests that the deformation started in the cretaceous. however, according to the apatite fission-track analysis the sedimentary rocks in the phu quoc basin and the kampot fold belt were buried to a depth where the temperature exceeded 100°c (fyhn et al. 2010). this indicates burial below 2–3 km of sediments that were entirely eroded away in the early eocene when the main exhumation 0 500 1000 1500 2000 2500 3000 3500 0 500 1000 1500 2000 2500 3000 3500 fwr fwr fwr lower cretaceous deposits neogene unconformity hwr hwr hwr neogene marine sediments pre-cretaceous deposits c. 1 km 1 km tw o -w ay t ra ve l ti m e (m se c) tw o -w ay t ra ve l ti m e (m se c) lower décollement ww ew efig. 3. seismic section representing the intermediate part of the tectonic complex. this part is characterised by increasing numbers of hangingwall anticlines and thrust faults. the upper thrust fault constitutes an upper flat to the west, an upper footwall ramp (fwr) that connects the upper flat to an intermediate flat, and a lower footwall ramp that continues into the lower flat to the east. each footwall ramp corresponds to a hanging-wall anticline above a displaced hangingwall ramp (hwr). due to the presence of two footwall ramps a prominent syncline developed between the two hanging-wall anticlines. blue lines: prominent bedding surfaces. red lines: thrust faults. purple line: the main neogene unconformity that truncates the structures in the tectonic complex. for location see fig. 1. 102102 of the region occurred (fyhn et al. 2010). it is suggested that the compressional deformation was caused by subduction of the westernmost part of the pacific ocean plate (metcalfe 1996). erosion of the up-thrust, pre-cretaceous rock units in the kampot fold belt supplied sediments for the cretaceous deposits in the phu quoc basin. the progressing compression led to deformation of the proximal part of the basin. a crosssection of the tectonic complex indicates a 200 km wide zone and the shortening of the complex is roughly estimated to be 50%. assuming a compressional orogenic translation in the order of 5 cm/y, the deformation lasted c. 4 million years. following the deformation, granites were intruded into the fold belt as the subducted sedimentary rocks below the fold belt began to melt. the age of the intrusions corresponds with the cretaceous deformation in the distal part of the tectonic complex. the crucial issue is to understand the mechanism that caused uplift in the eocene and subsidence in the neogene. we suggest that the subducted slab with a top layer of granitic composition was less dense than the base of the overlying lithosphere, and that this could cause a regional but orogenically passive uplift. the uplift was followed by eocene–miocene erosion and denudation before the raised granitic lithosphere cooled, which led to the neogene subsidence that created new accommodation space in the area above the phu quoc basin. conclusion the deposits in the phu quoc basin off south-west vietnam were affected by a late mesozoic orogeny. the tectonic complex at nam du archipelago is characterised by thin-skinned thrust-fault deformation with a distal part to the west, an intermediate part around the nam du archipelago, and a proximal part, which includes a hilly area in the onshore part of the kampot fold belt. the hinterland of the tectonic complex includes a magmatic arc represented by granitic plutons exposed in the south-western part of vietnam. the orogenic deformation of the complex is interpreted to be of early–middle cretaceous age. the complex was successively buried by a more than two kilometres thick package of sediments that was removed by erosion during uplift in the early eocene. finally subsidence in the neogene resulted in sedimentation that covered the mid-tertiary unconformity. acknowledgements we thank vietnam petroleum institute and petrovietnam for permission to use the commercial seismic data and permission to publish figs 2 and 3. the company landmark is thanked for a university grant to the department of geography and geology at copenhagen university. the enreca programme of danida is thanked for financial support. references fyhn, m., pedersen, s.a.s., boldreel, l.o., nielsen, l.h., green, f.p., dien, p.t., huyen, l.t. & frei, d. 2010: palaeocene – early eocene inversion of the phuquoc-kampot som basin: se asian deformation associated with the suturing of luconia. journal of the geological society (london) 167, 281–295. metcalfe, i. 1996. pre-cretaceous evolution of se asian terranes. in: hall, r., blundell, d. (eds.): tectonic evolution of southeast asia. geological society special publication (london) 106, 97–122. pedersen, s.a.s.2006: strukturer og dynamisk udvikling af rubjerg knude glacialtektoniske kompleks, vendsyssel, danmark. geologisk tidsskrift 2006, 1, 46 pp. pedersen, s.a.s., fyhn, m. dien, p.t., boldreel, l.o., nielsen, l.h., green, f.p., huyen, l.t. & mai, l.c. 2009: structural geology of the nam du island and neighbouring areas in the phu quoc basin, sw vietnam. danmarks og grønlands geologiske undersøgelse rapport 2009/7, 47 pp. authors’ addresses s.a.s.p. & l.h.n., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: sasp@geus.dk l.o.b., e.b.m. & m.b.f., department of geography and geology, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. geological survey of denmark and greenland bulletin 15, 2008, 21-24 the danish term ‘moler’ is the name for a special and unique marine deposit of lower eocene age found in the northern part of denmark and the danish north sea. in the literature it is often referred to as mo-clay, the english translation of ‘moler’ – a whitish, powdery sediment that lithologically is a clayey diatomite. the deposit, which is defined as the fur formation, is also well known for its 180 volcanic ash beds, increasing in number towards the top of the formation (pedersen & surlyk 1983). due to pleistocene glaciotectonic deformations the diatomite deposits crop out at the surface in the limfjorden area (gry 1940; klint & pedersen 1995; pedersen 1996, 2000). prior to the deformations the fur formation was situated at about 50–100 m below sea level, but during the deformations the diatomite was displaced upwards into glaciotectonic complexes. the complexes form elongate parallel hills up to 80 m a.s.l. in the western lim fjord region (fig. 1). the clayey diatomite attracts attention because it is a valuable raw material for production of insulation bricks and absorbing granulates, which are mainly used as cat litter. in addition, the exposed fur formation is a unique reference for investigations of the palaeogene stratigraphy in the north sea, where mudstones and shales with ash layers are known as the sele and balder formations (schiøler et al. 2007). in a tectonic framework the ash layers provide a unique addition to the understanding of the development of the north atlantic igneous province at the time when greenland and norway began to drift away from each other (larsen et al. 2003). moreover, the fur formation is especially noted for its rich fossil fauna, which comprises remarkably well-preserved specimens of birds, fish and insects. due to the public interest two museum exhibitions have been established, on mors and fur, and the geological features are so evident that numerous geological field trips have benefited from the success of well-displayed geology seen in exposures along the coastal cliffs and in the mo-clay pits. at present a norwegian drilling company is planning to use the fur formation outcrops at skarrehage for testing before applying their new drilling method offshore. in 2007, geus has continued many years of mo-clay investigations, and this paper presents some results from the 2007 activities, in addition to a review on the geology of the mo-clay. sedimentology of the fur formation the fur formation is c. 60 m thick at the type locality in the north-west corner of the island of fur. the sedimentology of the formation was described by pedersen (1981) and pedersen & surlyk (1983), who demonstrated that parallel well-laminated intervals alternate with structureless intervals. due to the exact identification and numbering of the ash layers the two sediment types can be followed throughout the dia tomite basin. the laminated mo-clay formed during anoxic periods, whereas the structureless mo-clay represents oxic events, during which a rich bottom fauna inhabited the seabed. in many layers the ash grains fill the burrows facilitating the recognition of the trace fossils (pedersen 1981). petrologically, diatom opal frustules make up 65% of the sediment. the dominant diatom is coscinodiscus (fig. 2). the clay content is about 28%, but varies, and the clay fraction comprises the clay mineral smectite (pedersen et al. 2004); the remainder is mainly disseminated pyrite. mo-clay situated above the groundwater level has been depleted of pyrite, but mo-clay situated below the groundwater level shows a pyrite content that varies from 5 to 8%. this gives the diatomite a black colour. above the redox level the diatomite is light coloured due to the depletion of pyrite, which is corroded and transformed into sulphates (jarosite and gypsum) palaeogene diatomite deposits in denmark: geological investigations and applied aspects stig a. schack pedersen © geus, 2008. geological survey of denmark and greenland bulletin 15, 21–24. available at: www.geus.dk/publications/bull 21 fig. 1. map of the western limfjorden region with place names. note the hilly landscapes on the islands of mors and fur. 10 km thy mors salling fur limfjorden thisted siltstrup stærhøj hanklit skarrehage feggeklit ejerslev fur knudeklint nykøbing ertebølle junget n 57°n 57°n 9°e 9°e denmark and limonite. part of the latter is transported out to the slopes and cliffs at the coast, where it cements the quaternary deposits, mainly the glaciofluvial sand and gravel, to form the so called ‘red stone’. during diagenesis amorphous opal a is transformed to microcrystalline opal c/t, and in the lower part of the formation cherty shales (skiferlag) have formed (pedersen et al. 1998). in the middle and upper part of the formation calcareous concretions are seen. the concretions are found in about five marker levels, they are ellipsoid in shape and their size range from 0.3 to 0.75 m in thickness and one to a few metres in lateral extent. one exception is the calcareous concretion cementing the ash layer +101, which is a 35 cm thick marker bed in the upper part of the formation. gypsum is sometimes found as small rosettes up to 5 cm in size, often concentrated in glaciotectonic shear zones. volcanic ash layers the volcanic ash layers were stratigraphically logged and numbered by bøggild (1918), who also published the first data on the petrological composition and is responsible for the division of the ash layers into a negative and a positive series. from their geochemical characteristics, four stages of volcanic activity can be recognised (pedersen et al. 1975). stage 1 corresponds to the ash layers from –39 to –22, which vary in composition from rhyolitic to basaltic. the source of these ash layers was volcanic centres situated on the shelf west of the british isles (larsen et al. 2003). a volcanic stage pre-dating stage 1 is recorded from offshore britain (knox et al. 1997) and is probably responsible for the bentonite in the holmehus formation; it is well known that bentonite is a devitrification product of volcanic glass. stage 2 is represented by ash layers –21 to –15, which are variable in composition and comprise phonolites, nepheli nites, trachytes and rhyolites. the peralkaline nephelinitic ash layer –19 is a remarkable blue ash layer, easily recognisable, and the very distinct, 4 cm thick, orange coloured –17 is a diagenitically altered trachyte, formerly thought to be rhyolitic. the source of the stage 2 ash layers was either the shelf area west of the british isles or the gardiner igneous complex in east greenland (larsen et al. 2003). stage 3 is represented by the three distinctive black alkali basalt ash layers –13, –12 and –11. the ash layers may originate from the opening rift between north-west europe and greenland (larsen et al. 2003). stage 4 is the main tholeiitic basalt eruption, during which the positive numbered ash layers were formed. they correspond in composition to the basalts currently forming in iceland. thus the source could be regarded as a proto-iceland situated in the middle of the incipient oceanic crust of the north atlantic ocean (larsen et al. 2003). the most significant ash layer in this unit is the 20 cm grey rhyolitic andesite +19, which has been dated to 54 ma (larsen et al. 2003), and which is an important marker bed for the exploitation activities. glaciotectonic deformation in the 1930s state-geologist h. gry consulted the mo-clay companies, and from his co-operation with them he realised how important structural geology was for solving the exploration problems. in 1940 he summarised the structural knowledge of all the known outcrops of the fur formation and described their glaciotectonic framework (gry 1940). gry’s model was monoglacial with a lobe-shaped body of ice advancing from the north. pedersen (2000) demonstrated that the glaciotectonic structures were due to superimposed deformation. the norwegian and swedish advances spreading over the northern part of denmark in the weichselian at about 28–24 ka b.p. created the main structural features. however, evidence of older glaciations of saalian age is also recorded, both as deposits (hesselbjerg till and harhøj sand series) and as minor deformations along the glaciotectonic unconformity truncating the fur formation (fig. 3). the advanced structural studies of the glaciotectonic complexes are based on the concept of thin-skinned thrust-fault 22 fig. 2. scanning electron microscope images of clayey diatomite from the fur formation. the upper image is a plan view, where the large diatom frustules of coscinodiscus are seen. the lower image is perpendicular to the lamination and shows some diatom frustules in cross-section. 200 µm a 200 µm b deformation and construction of balanced cross-sections. a prominent example of this is the hanklit glaciotectonic complex, which constitutes three 60 m thick thrust sheets with a maximum displacement of about 300 m from the décollement surface at 80–100 m below sea level (klint & pedersen 1995). before the displacement the upper surface of the fur formation was situated 20 m below sea level, from where it was displaced to 60 m above sea level to form the elongated e–w-trending hills of northern mors (fig. 1). a similar depth to the décollement surface (about 100 m b.s.l.) was calculated from area balance of the feggeklit cross-section (fig. 1). this study also demonstrated that the sequential deformations within a glaciotectonic propagation were related to one ice advance (pedersen 1996). a number of advanced glaciotectonic studies are described in the series danmarks og grønlands geologiske undersøgelse rapport from 1996 to 2007, and document the structural framework of the areas planned for future excavation of the clayey diatomite. raw material investigation in general all the mo-clay pits are located within elongate hanging-wall anticlines in thrust sheets of the fur formation. thus the basic problem is to identify the trend of the anticlines in order to locate the boundaries of the excavation areas. when this has been solved, the important two questions to be answered are: how much diatomite is present, and how much cover has to be removed during the excavation? the answers are provided by construction of isopach maps of the excavation areas and volume calculation with application of kriging (cressie 1990). the data applied for the kriging are based on the information from the exploration wells (fig. 4), where the depth to ash layer +19 is the most important parameter. the ash layers above +19 are included in the 23 a b -303 -304 -305 0–2 2–4 4–6 6–8 8–10 10–12 12–14 14–16 16–18 18–20 >20 calculation area borders of private property investigation wells thickness of quaternary cover, september 2005 (m) 0–2 2–4 4–6 6–8 8–10 10–12 12–14 14–16 16–18 18–20 >20 calculation area borders of private property investigation wells thickness of quaternary cover, september 2005 (m) 100 m 100 m calculation area fig. 4. two isopach maps constructed by kriging calculation based on well data. a: gives an unrealistic ne-trend of the raw material deposit. b: improves the reliability of the isopach map by including data from the three new wells, –303, –304 and –305, in the calculation. this map agrees well with the original structural model for the area. 10 m w dgu well no 31.304 dgu well no 31.303 topographic surface ejerslev mo-clay pit excavation surface in pit ash layer 101 ash layer 19ash layer 1 e ash-layer series upper mo-clay series calcereous concretion black mo-clay ejerslev till fegge till harhøj sand beds hesselbjerg till pre-quaternary unco nform ity fig. 3. geological cross-section of the area with black diatomite. the location of the wells is shown in fig. 4b (–303 and –304). the hesselbjerg till and the harhøj sand beds are of saalian age, whereas the fegge and ejerslev tills represent, respectively, the weichselian norwegian and swedish advances. 24 unutilised cover, which is dominated by quaternary deposits (fig. 3). only few and thin ash layers are found below ash layer +19, and they do not affect the production. the lower boundary of the excavation is either at the groundwater level or at the occurrence of black diatomite, which usually coincide. however, exceptions occur as illustrated below. the ejerslev field is currently the main production area on mors. for the production planning, an isopach map (fig. 4a) was constructed in the spring of 2005. however, the density of data points is crucial for the calculation of the exploitation area, and the resulting map contradicted the structural model for the area (fig. 5). in the spring of 2007 an unusual occurrence of black diatomite cropped out during excavation of the uppermost part of the fur formation (fig. 6). based on three new wells, combined with field investigations, it was concluded that the black diatomite is preserved in an anticlinal structure (fig. 3). since the late weichselian (c. 20 ka) this dome feature had been protected from percolating groundwater and no leaching of the pyrite had taken place. the subsequent kriging analysis had an important new data point in the most poorly covered area. the new isopach map (fig. 4b) supports the original structural model for the area (fig. 5) demonstrating the importance of using structural models. references bøggild, o.b. 1918: den vulkanske aske i moleret. danmarks geologiske undersøgelse ii. række 33, 84 pp. cressie, n.a.c. 1990: the origins of kriging. mathematical geology 22, 239–252. gry, h. 1940: de istektoniske forhold i molerområdet. meddelelser fra dansk geologisk forening 9, 586–627. klint, k.e.s. & pedersen, s.a.s. 1995: the hanklit glaciotectonic thrust fault complex, mors, denmark. danmarks geologiske undersøgelse serie a 35, 30 pp. knox, r.w.o’b., holloway, s., kirby, g.a. & bailey, h.e. 1997: strati graphic nomenclature of the uk north west margin: 2. early pale ogene lithostratigraphy and sequence stratigraphy, 58 pp. notting ham: british geological survey. larsen, l.m., godfrey fitton, j. & pedersen, a.k. 2003: paleogene volcanic ash layers in the danish basin: compositions and source areas in the north atlantic igneous province. lithos 71, 47–80. pedersen, a.k., engell, j. & rønsbo, j.g. 1975: early tertiary volcanism in the skagerrak: new chemical evidence from ash layers in the moclay of northern denmark. lithos 8, 255–268. pedersen, g.k. 1981: anoxic events during sedimentation of a palae ogene diatomite in denmark. sedimentology 28, 487–504. pedersen, g.k. & surlyk, f. 1983: the fur formation, a late paleocene ash-bearing diatomite from northern denmark. bulletin of the geological society of denmark 32, 43–65. pedersen, g.k., pedersen s.a.s., steffensen, j. & pedersen, c.s. 2004: clay content of a clayey diatomite, the early eocene fur formation, denmark. bulletin of the geological society of denmark 51, 159–177. pedersen, s.a.s. 1996: progressive glaciotectonic deformation in weich selian and palaeogene deposits at feggeklit, northern denmark. bulletin of the geological society of denmark 42, 153–174. pedersen, s.a.s. 2000: superimposed deformation in glaciotectonics. bulletin of the geological society of denmark 46, 125–144. pedersen, s.a.s., lindgreen, h. & pedersen, g.k. 1998: amorphous silica and hydrous aluminosilicates for production of construction materials. inco-copernicus project no. erbic15ct96 0712. danmarks og grønlands geologiske undersøgelse rapport 1998/3, 47 pp. poul schiøler et al. 2007: lithostratigraphy of the palaeogene – lower neogene succession of the danish north sea. geological survey of denmark and greenland bulletin 12, 77 pp. author’s address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark, e-mail: sasp@geus.dk fig. 6. drill operation in the ejerslev clay pit. the vehicles are parked on the top of the barkær anticline (see fig. 5), which is seen in front of the drill rig. note the black colour of the diatomite, which is caused by a high content of disseminated pyrite. till deposits occur behind and above the drill rig. compare with fig. 3. thrust fault barkær anticline glacial sediments ash-layer series upper diatomite series lower diatomite series fig. 5. structural model for the northern part of the ejerslev field. the barkær anticline is the main target for exploitation. in the syncline east of the barkær anticline, a more than 20 m thick sequence of quaternary deposits is found, which makes this area uneconomic for excavation. moreover, the diatomite here is probably also of a black variety unsuitable for production. geological survey of denmark and greenland bulletin 20, 2010, 75–78 75 establishment of robust reservoir models and estimates of subsurface hydrocarbon volumes in relatively unknown subsurface settings can be improved by using data from field analogues. the discovery of the rosebank oilfield in the faroe–shetland basin showed that intrabasaltic sandstones can form important hydrocarbon reservoirs in volcanic basins (helland-hansen 2009). the sødalen region in southern east greenland (fig. 1) forms an excellent field analogue to the rosebank oilfield where contemporaneous palaeogene sediments interbedded with lava units can be studied and sampled (larsen et al. 1999). in this area many of the exposures are located along steep, inaccessible cliffs with excellent exposures that are ideal for 3-d photogeological studies based on digital high-resolution photographs taken from a helicopter. the analogue study reported on here has integrated results from a wide range of spatial scales. on a large scale (kilometre to metre), 3-d photogeology was used to study the extent, geometry and interfingering of volcanic and intrabasaltic sedimentary units. photogeology was also used to map faults, dykes and sills, which may lead to compartmentalisation (division) of reservoirs. on an intermediate scale (metre to millimetre), sections are logged in the field, sedimentary and volcanic facies are mapped and depositional environments are interpreted. three-dimensional photogeology is also applied on an intermediate scale to map lateral variations of sedimentary units between logged sections. on a small scale (millimetre to micrometre), mineral-chemical, petrographical and zircon age determinations provide information on sediment source, provenance area and diagenetic influences on reservoir properties. the analogue study has resulted in a large database, which can form an important source of estimates of reservoir size, geometry and connectivity, and of vertical and lateral variations in the sandstone content of reservoirs. ultimately this may improve estimates of the actual volumes and recoverable volumes of hydrocarbons in intrabasaltic subsurface sediments. the 3-d photogeological method the 3-d photogeological method used in this study was developed at the geological survey of denmark and greenland (geus) and builds on earlier work by dueholm et al. (1993) and dueholm & olsen (1993). the method allows the acquisition of geological data from vertical and oblique aerial photographs, with a three-dimensional overview of the outcrops. the oblique photographs (1:15  000–1:17 000 scale) are triangulated with coloured, vertical, aerial photographs (1:27 000 scale) using a 3-d stereo-plotter coupled with stereo-mirror technology. the mapping of geological features includes determination of strata thickness, strike direction and dip values working on a 3-d high resolution vision of the cliffs. the resolution of volcanic and sedimentary beds and geological features is c. 10 cm. all the mapped features are stored in a gis database and 3-d polylines can be exported study of a palaeogene intrabasaltic sedimentary unit in southern east greenland: from 3-d photogeology to micropetrography henrik vosgerau, pierpaolo guarnieri, rikke weibel, michael larsen, cliona dennehy, erik v. sørensen and christian knudsen sø dalen ‘sødalengletscher’ 68°15´n 31°w miki fjord j.c. jacobsen fjord greenland mainly palaeogene volcanic rocks palaeogene gabbro mesozoic–palaeogene sedimentary rocks precambrian basement 10 km fig. 2a fig. 1. map of the sødalen area in southern east greenland. the red line shows the location of the profile in fig. 2a. © geus, 2010. geological survey of denmark and greenland bulletin 20, 75–78. open access: www.geus.dk/publications/bull 7676 a. large scale b. intermediate scale c. small scale 0 1 2 3 4 5 6 7 lava unit invasive lava lower–middle shoreface lower–middle shoreface lava unit 8 9 10 m cl si sand gnl 5 m 20%10090%80706050403020100 1 1 2 2 upper shoreface dyke 20 μm b volcanics volcanics chlorite quartz feldspar volcanic fragments mica heavy minerals sphene rock fragments quartz feldspar calcite illite chlorite detrital composition authigenic phases b 200 m bedding dyke fault intrabasaltic unit nwnw sesenw se 77 as shape files suitable for 3-d modelling using, for example, petrel reservoir engineering software. moreover, using 3-d feature databases in arcgis, geological cross-sections can be generated automatically to obtain real representations of outcrops, and then projected onto a topographic profile, where the accuracy is as high as the resolution in the photographs. the oblique photographs used here were small-frame colour photographs taken from a helicopter flying close to the cliff faces (<800 m) and at a constant altitude along straight lines approximately parallel to the cliffs. the photographs were taken with a 60 to 80% overlap using a 22 megapixel digital camera. on a large scale (kilometre to metre) on a large scale, 3-d photogeology is used to study the lateral extent and geometry of the intrabasaltic sediments and volcanic rocks and boundary relationships. evidence of compartmentalisation of the intrabasaltic reservoir analogues, caused for example by dykes, sills or faults, are mapped. figure 2a shows an oblique view of a 1.2 km section on the eastern side of sødalen, which is an 8 km long u-shaped valley, orien tated se–nw from miki fjord to ‘sødalengletscher’ (fig. 1). the photograph focuses on the stratigraphically lowest intrabasaltic, whitish sedimentary unit, which dips gently to the south-east. the geological cross-section in fig. 2a is projected on the cliff view, and is obtained from 3-d polylines created during the 3-d photogeological work. figure 2a illustrates several large-scale features relevant to the analogue study such as: (1) top and bottom geometry of the sandstone unit, (2) density of dykes and faults, which has a large impact on the lateral extension of the layers due to offsetting and (3) an evaluation of reservoir compartmentalisation. on an intermediate scale (metre to millimetre) on an intermediate scale, sedimentary and volcanic sections are logged in the field. facies types are identified, their lateral distribution and vertical stacking patterns are mapped and the boundaries between sedimentary and volcanic units are studied in detail. the 3-d photogeology is also useful on this intermediate scale because the high resolution of the digital photographs allows enlargement to study decimetre-sized features. photogeology can therefore be very helpful in mapping sedimentary facies assemblages between logged sections as well as key surfaces separating the different facies, such as sequence boundaries and marine-flooding surfaces. the overall depositional environments and the governing mechanisms for facies distribution, such as sediment transport directions, relative sea-level variations and palaeotopography can be interpreted from these studies. an example of a study on an intermediate scale is illustrated in fig. 2b where the photograph is an enlargement of a small area on the digital photograph (fig. 2a). the log of the sedimentary unit of shallow marine sandstone is shown on the left side of the figure. the lower, exposed part of the unit consists of crudely and irregularly bedded sandstones, locally with vertical burrows, interpreted as deposited in the upper shoreface zone. the lower part is overlain by wellsorted, fine to medium-grained, laterally extensive sandstone sheets and wedge-shaped sandstone beds with a large variety of sedimentary structures including local vertical burrows, cross-stratification and parallel bedding. these sandstones are interpreted as deposited in the lower to middle shoreface zone, which implies that the boundary to the underlying upper shoreface sandstones represents a minor flooding surface. on the photograph (fig. 2b) it is seen that the upper and lower to middle shoreface facies assemblages can be followed laterally for tens of metres. it is also seen that the boundary between the lava units and the sedimentary unit is slightly undulating, and that the invasive lava bed can be followed into the overlying lava to the right. the dyke that cuts through both the sedimentary unit and the lava units may have led to a possible compartmentalisation of the sandstone reservoir. fig. 2. (facing page) a: oblique photograph (upper) and derived geological cross-section (lower) of the eastern side of sødalen, an 8 km long u-shaped glacial valley, extending se–nw from ‘sødalengletscher’ to miki fjord. for location see fig.1. the photograph focuses on the lowest intrabasaltic sediments (whitish colour) that dip gently to the south-east. faults and dykes are also seen. the geological cross-section below is a projection of 3-d polylines along a n–s-oriented profile, slightly different from the nw–se orientation of the photograph. b: close-up view of the white square in fig. 2a. a field log of the section (left) labels the sedimentary facies assemblages which can be followed laterally on the photograph. other important observations include an invasive lava bed and a dyke cutting through both lavas and intrabasaltic sediments. c: detrital and authigenic mineralogical composition of the facies in the sedimentary unit. the larger amount of authigenic phases in the lower-middle shoreface sandstones compared to the upper shoreface sandstones, reflect that these sandstones originally had a larger content of unstable, glass-rich volcanic fragments. the positions of the two samples are shown on fig. 2b. 7878 on a small scale (millimetre to micrometre) on a small scale, petrography is used to understand diagenetically induced reductions in porosity and permeability to understand the influence of provenance, sedimentary facies and surrounding ‘hot’ units on diagenetic processes. based on intensive sampling in a well-described geological framework controlled by 3-d photogeology and logged sedimentary sections, diagenetic changes are compared with detrital composition, depositional environment and effects from overlying and underlying lava units as well as dykes and sills. provenance variations are revealed from heavy mineral analysis using computer-controlled scanning electron microscopy, zircon age distributions and petrography. geochemistry is applied to distinguish different intrabasaltic units. the intrabasaltic sedimentary rocks consist of a mixture of detrital siliciclastic (quartz, feldspar, mica etc.) and volcaniclastic input (fig. 2c). the upper shoreface facies is richer in volcanic fragments than the lower-middle shoreface facies, yet it has a lower content of authigenic phases (fig. 2c). this is unexpected as volcanic fragments traditionally have been associated with intensive alteration thereby liberating elements for extensive authigenic phases. however, the type of volcanic fragments is also crucial for the degree of diagenetic alteration. glass-rich volcanic fragments are common in the lower-middle shoreface facies, whereas relatively stable volcanic fragments (lath-shaped plagioclase with little interstitial glass matrix) are more abundant in the upper shoreface. glass-rich volcanic fragments, which are easily altered, result in extensive authigenic formation, including chlorite as shown in fig. 2c. the stable volcanic fragments behave as plagioclase grains during diagenesis and have less influence on the authigenic phases than the glass-rich volcanic fragments. consequently, the upper shoreface sandstones show better reservoir properties than the lower-middle shoreface facies. conclusions the field analogue project at sødalen integrated the three disciplines of 3-d photogeology, sedimentology and petrography, and gave detailed information from kilometre to micrometre scale. petrographical investigations revealed the diagenetic influence on the reservoir properties. when the diagenetic changes were related to the sedimentary facies, the information on the reservoir properties could be scaled up to sedimentary bodies. the geometry of the sedimentary bodies and the probability of compartmentalisation are defined from 3-d photogeology and logged sedimentary sections. integration and up-scaling of several types of geological data resulted in a more complete understanding of the geology of the area and can form the basic input for reservoir modelling and as field analogue for hydrocarbon discoveries in a similar, inaccessible geological setting in offshore areas. acknowledgements chevron and sindri group are thanked for their financial contribution to the field work in greenland. references dueholm, k.s. & olsen, t. 1993: reservoir analog studies using multimodel photogrammetry; a new tool for the petroleum industry. aapg bulletin 77, 2023-2031. tulsa, oklahoma: american association of petroleum geologists. dueholm, k.s., garde, a.a. & pedersen, a.k. 1993: preparation of accurate geological and structural maps, cross-sections and block diagrams from colour slides, using multi-model photogrammetry. journal of structural geology 15, 933–937. helland-hansen, d. 2009: rosebank – challenges to development from a subsurface perspective. in: varming, t. & ziska, h. (eds): faroe islands exploration conference: proceedings of the 2nd conference. annales societatis scientarium faeroensis, supplementum 50, 241–245. larsen, m., hamberg, l., olaussen, s., nørgaard-pedersen, n. & stemmerik, l. 1999: basin evolution in southern east greenland; an outcrop analog for the cretaceous–paleogene basins on the north atlantic volcanic margin. aapg bulletin 83, 1236–1261. tulsa, oklahoma: american association of petroleum geologists. authors’ addresses h.v., p.g., r.w., e.v.s. & c.k., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: hv@geus.dk m.l., dong energy a/s, agern alle 24–26, dk-2970 hørsholm, denmark. c.d., chevron upsteams europe, chevron north sea ltd., chevron house, hill of rubislaw, aberdeen ab15 6xl, uk. geological survey of denmark and greenland bulletin 15, 2008, 81-84 the water framework directive (wfd) of the european union aims to achieve a ‘good’ status for all inland and coastal waters by the year 2015 (ec 2000). the directive defines how this should be achieved through the establishment of environmental objectives and ecological targets. successful implementation of the wfd requires integration into already existing national legislation and a sound combination of issues on technical feasibility, scientific knowledge and socio-economic aspects requiring intensive stakeholder involvement. this calls for appropriate tools such as models to support management of technical and social aspects of different phases of the implementation (rekolainen et al. 2003; quevauviller et al. 2005). it is therefore necessary to provide an overview of already existing methods and tools and develop new ones. research programmes funded by the european commission (ec) often address issues of current interest for practitioners, such as the fifth framework programme, where a number of research projects to support the practical implementation of the wfd were initiated under the theme ‘energy, environment and sustainable development’. the funding part (the directorate-general for research, dg research) and the responsible authority for the wfd at european level (directorate-general of environ ment) saw the need to cluster these research projects and related activities, and initiated the harmoni-ca project, a socalled ‘concerted action’ (i.e. harmonised modelling tools for integrated river basin management). the objectives of this paper are (a) to briefly describe the overall purpose of the harmoni-ca project and some of its overarching outputs, and (b) to further illustrate how the implementation of the wfd can be enhanced by combining monitoring and modelling disciplines and by bringing practitioners and researchers together. harmoni-ca the harmoni-ca project started in october 2002 and concluded with a major conference in brussels in september 2007. the main objectives of the project were: (1) to build a bridge from research to practice; (2) to create a forum for related research projects to exchange ideas, to optimise and co-ordinate activities in ongoing research projects, and to initiate new spin-off projects; and (3) to gather already existing information, experience and research on both national and european levels that can support the implementation of the wfd (arnold et al. 2005). bridging research and practice a central activity of the project was to bring practitioners and researchers together at a large number of targeted workshops and open annual conferences. the main purpose was to open the floor for discussions on the needs of those working with integrated water management and the related outputs from the scientific community. several opportunities and new ideas arose from the workshops. however, obstacles and bottle necks were also recognised, such as insufficient dialogue be tween the scientific and policy-making communities due to different interests and languages, lack of translation of scientific outputs into tools readily applicable to policy-makers, and the lack of a structure within which the groups respon sible for the implementation of the wfd could be brought together with the scientists (arnold et al. 2005). while the open annual conferences had broad themes, the workshops concentrated on the following tasks of harmonica: (1) establishment of a ‘tool box’ to provide easy and guided access to information and communication technologies for the development of river basin management plans (van griensven & vanrolleghem 2006); (2) development of a generalised methodological framework for harmonised model support in integrated river basin management (hat termann & kundzewicz 2006); (3) better integration be tween monitoring and modelling in water management (højberg et al. 2007a, jørgensen et al. 2007); and (4) investigation of how the science–policy interface can be bridged in current water management (borowski & hare 2007). more than 20 targeted workshops addressing these issues were held, resulting in workshop reports, synthesis reports and har moni-ca guidances. forum for research projects the activities in harmoni-ca and other related research projects were integrated in a cluster called catchmod. the projects in this cluster produced outputs that could support the implementation of the wfd in different ways. this co-ordi81 from science to practice in implementing the european union’s water framework directive lisbeth flindt jørgensen, jens christian refsgaard and anker lajer højberg © geus, 2008. geological survey of denmark and greenland bulletin 15, 81–84. available at: www.geus.dk/publications/bull nation of research and technology development activities supported researchers in exchanging ideas on modelling tools to support the wfd. it aimed to increase the output and benefit of ongoing research, to speed up the (re-)use of developed products, to avoid major overlaps between projects,and to reduce the risk of duplicating activities. harmoni-ca achieved this by, amongst other things, organising two technical catchmod workshops for the projects involved, where a dialogue was established among the scientific communities and gaps were reduced between different research disciplines. the workshops were also often used as an instrument to prepare input to guidance documents and synthesis reports. support to the implementation of the wfd another important task of the project was to collect already existing knowledge on issues related to the different steps of the wfd. this was partly done by producing a number of reports and guidances and partly by establishing a web portal. synthesis reports and guidances. a large number of synthesis reports and guidances was initiated by harmoni-ca. the following groups can be distinguished: (1) reports supporting modelling activities and development, e.g. quality assurance, sensitivity analysis, and decision support development; (2) reports supporting the collaboration between different scientific/policy fields such as monitoring, modelling, agriculture, economy and (3) reports on the science–policy interface and end-user involvement. the reports harmonise available knowledge, provide added value by picking up recent insights and are essential in order to improve the quality of communication between science and practice. most of the reports have been commissioned to small task groups consisting of both scientists and end-users and include outcomes of discussions held at various workshops and conferences. some reports will be published in the international water association publication series, improving the visibility of european research. examples are: ‘uncertainty analysis’ (refsgaard et al. 2005), ‘model-supported implementation of the water framework directive – a water manager’s guide’ (hattermann & kundzewicz 2007), ‘integration of the human dimension in model-supported water management’ (bots et al. 2007) and ‘good practice in joint use of monitoring and modelling’ (højberg et al. 2007b). the wise-rtd web portal. it is often a difficult task for practitioners dealing with new or challenging steps of the wfd, to find their way through the jungle of existing knowledge and experiences, when they look for assistance or good examples. harmoni-ca has therefore developed a web portal (www.wise-rtd.info) that mainly acts as an entry to tools, experiences, guidances and research activities or results that can help water managers or others interested in finding information relevant to the wfd implementation. there are several ways to search for help in the portal. users can enter in their capacity as water managers, scientists or stakeholders. another option is to search by keywords de rived from the ‘common implementation strategy guid ance documents’ developed by the ec to support the implementation of the wfd. information on tests or pilot projects can also be found, such as the ‘pilot river basins’, where different steps of the wfd have been tested before final implementation. the portal is hosted and supported by the euro pean commission and is expected to become the support portal for wfd implementation. joint use of monitoring and modelling while a combination of monitoring and modelling is often seen in research, there seems to be more hesitance to use modelling in practical water management where, on the other hand, a lot of data acquisition takes place. one of the tasks within harmoni-ca was to try to integrate the monitoring and modelling disciplines in water management to a higher degree than currently seen. for this purpose five workshops have been arranged over a period of three years. a total of more than 80 water managers, stakeholders, consultants, policy makers and scientists participated, representing 24 mainly european countries. the first three workshops investigated the status of monitoring programmes and the present use or knowledge of modelling support to monitoring. it was recognised that monitoring programmes often date back several decades and have traditionally been considered an independent discipline. however, within 82 500 km campania region ezousas aquifer ialomita river basin nevesis river basin pärnu river basin hørup well field fig. 1. map of europe showing the locations of the six case studies discussed. the last few decades modelling has entered the arena as a supplementary tool to help extract information from observation data. this is generally accepted in the research community. however, in practical water management there is considerable reluctance to employ models due to various ob stacles such as lack of skill, lack of time, lack of awareness on what models can do and also lack of confidence in models (brugnach et al. 2007). it was therefore decided to use the last two workshops to discuss six case studies from different areas of europe (fig. 1), in order to explore the possibility of developing an outline for a common approach in implementing the wfd and combining monitoring and modelling activities. these case studies are briefly presented here: hørup well field (denmark). groundwater extraction for drinking water results in low base flows in nearby streams in dry seasons. there are threats to the groundwater quality from diffuse pesticide pollution and from point sources (contaminated sites in a nearby city). the challenges in this case study were to ensure sustainable extraction without unacceptably affecting nearby streams and to protect the groundwater against pollution. ezousas aquifer (cyprus). heavy abstraction for irrigation causes saline intrusion into the aquifer. groundwater recharge is low due to a decline in precipitation and damming of the river that previously supplied most of the recharge. the challenges were to convince stakeholders of the positive effects of a planned artificial recharge programme using cleaned wastewater, to optimise this programme to avoid saline intrusion and to evaluate the present monitoring programme. pärnu river basin (estonia). threats to wetland and groundwater quality due to agricultural activities with both diffuse and point sources of pollution. peat mining in the area leads to local acidification of surface waters. the challenges were to differentiate diffuse and point source contamination in surface water and groundwater and to address the acidification problem. nevesis river basin (lithuania). surface waters are threatened by high nutrient loads from diffuse agricultural sources and from sewage from small villages without waste water treatment. the challenge was to differentiate between diffuse and point source contamination in surface water. campania (italy). the groundwater, and thus the drinking water, is polluted by agriculture and horticulture, and outlets of untreated waste water from small villages contribute to poor groundwater quality. the challenge was to differentiate between diffuse and point source pollution. ialomita river basin (romania). high levels of nutrients, especially nitrate, in both surface and groundwater, caused by aerial deposition from neighbouring areas with agricultural activities as well as by diffuse pollution from the local agriculture. in addition, a special problem arises from high nitrate levels in the groundwater due to extensive use of nutrients in the past. the challenges were to differentiate between contamination caused by aerial deposition, local sources, and the inherited high nitrate concentrations from earlier agricultural activities. each case study was elaborated by a small group of participants, and flowcharts were prepared showing how to implement the wfd, with special reference to combining models and monitoring. the groups working on the six cases fo cussed on different aspects. in spite of this, their flowcharts showed many similarities in their approaches, which allowed the construction of a common flowchart representing the key 83 modelling process understanding modelling scenario analyses cost-effectiveness modelling assess effects of pom modelling assess impacts modelling qa of data modelling inter/extrapolation modelling design of monitoring programmes surveillance monitoring problem indentification steps according to the wfd monitoring according to the wfd suggested modelling support investigative monitoring operational monitoring programme of measures need for more data? conceptual model baseline scenario no required confidence? rbmp implementation goals achieved? no yes st ak eh ol de r co ns ul ta tio ns an d in vo lv em en t yes fig. 2. flowchart for integrating monitoring and modelling activities when implementing the european union’s water framework directive. the chart is a result of workshops on six case studies from different parts of europe. the workshops addressed diverse challenges and problems of current interest. rmbp, river basin management plans; pom, programme of measures; qa, quality assurance. aspects of all the six cases (fig. 2). this flowchart may be applicable in most areas of europe that face different challenges in the implementation of the wfd. as shown in fig. 2, models can support a variety of different tasks. a numerical model can be used to test different conceptual models and to check whether these are consistent with all available data. a model can be used to evaluate the effects of already implemented or planned measures, and by undertaking analyses of different scenarios help to choose the best programme of future measures to improve the status of the environment. incorporation of data in models provides an additional check of the consistency and quality of the data, and since such data are often discrete in time and space, models are more applicable for interpolation and extrapolation than statistical methods. models can also help to achieve an optimal design of monitoring programmes by providing input on when and how often to monitor. they can also help to find a predefined confidence level, taking uncertainty into consideration. thus models can support data acquisition in many ways, and this support may go much further than the traditional prediction of effects of various alternative initiatives. although the six case studies were real-life situations addressing topics of current interest, a major limitation was the fact that the participants did not have to consider political or economic constraints. nevertheless, people with different backgrounds – scientists, practitioners, monitoring, modellers and experts – could easily work together in combining monitoring and modelling in an effective manner. all participants found the use of models to be beneficial in most of the situations mentioned above. concluding remarks different professional communities have different interests, traditions and working cultures. this constrains the uptake and acceptance of research results by practitioners and hinders the interaction of different disciplines such as monitoring and modelling. experience from harmoni-ca workshops on monitoring and modelling shows that when researchers and practitioners are brought together to elaborate on real life cases, they can easily work together in a very inspiring and constructive way. a key conclusion from the project is that it is possible to bridge some of the gaps between research and practice and between different disciplines, but that this re quires continuous attention and positive co-operation from all parties involved. acknowledgement the present work was carried out within the concerted action harmonica, which was funded under the european commission’s 5th framework programme (contract evk1-ct2001-00192). references arnold, g.e., de lange, w.j. & blind, m.w. 2005: the concerted action harmoni-ca: facilitating the dialogue and bridging the gap between research and wfd implementation. environmental science and policy 8, 213–218. borowski, i. & hare, m. 2007: exploring the gap between water managers and researchers: difficulties of model-based tools to support practical water management. water resources management 21, 1049–1074. bots, p., gooch, g., mcintosh, b.s. & pahl-wostl, c. 2007: integration of the human dimension in model supported water management. a harmoni-ca guiding document. available on: www.harmoni-ca.info. brugnach, m., tagg, a., keil f. & de lange, w.j. 2007: uncertainty matters: computer models in the science-policy interface. water resources management 21, 1075–1090. ec 2000: european commission, directive 2000/60/ec of the european parliament and of the council of october 23, 2000 establishing a framework for community action in the field of water policy. official journal of the european communities 2000, 22.12.2000, l327/1–l327/72. hattermann, f.f. & kundzewicz, z.w. 2006: model-supported integrated river basin management – methodological framework. in: kundze wicz, z.w. & f.f. hattermann (eds): natural systems and global change. poznan: rcafe pas, 65–74. hattermann, f.f. & kundzewicz, z.w. (eds) 2007: model-supported implementation of the water framework directive – a water manager’s guide. a harmoni-ca guideline. available on: www.harmonica.info. højberg, a.l., refsgaard, j.c., van geer, f., jørgensen, l.f. & zsuffa, i. 2007a: use of models to support the monitoring requirements in the water framework directive. water resources management 21, 1649–1672. højberg, a.l., refsgaard, j.c., jørgensen, l.f., van geer, f. & zsuffa, i. 2007b: good practice in joint use of monitoring and modelling. a catchment modelling guidance. available on: www.harmoni-ca.info. jørgensen, l.f., refsgaard, j.c. & højberg, a.l. 2007: the inadequacy of monitoring without modelling. journal of environmental monitoring 9, 931–942. quevauviller, p. et al. 2005: science-policy integration needs in support of the implementation of the eu water framework directive. environmental science & policy 8, 203–211. refsgaard, j.c., van der sluijs, j.p., højberg, a.l. & vanrolleghem, p.a. 2005: uncertainty analysis. a catchment modelling guidance. available on: www.harmoni-ca.info. rekolainen, s., kämäri, j. & hiltunen, m. 2003: a conceptual framework for identifying the need and role of models in the implementation of the water framework directive. international journal on river basin management 1, 347–352. van griensven, a. & vanrolleghem, p.a. 2006: the catchmod toolbox: easy and guided access to ict tools for water framework directive implementation. water science & technology 53, 285–292. 84 authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: lfj@geus.dk geological survey of denmark and greenland bulletin 20, 2010, 71–74 71 an airborne hyperspectral survey was organised by the geological survey of denmark and greenland (geus) and carried out in 2000 to test the use of spectral analysis in mineral exploration under arctic conditions. the hyperspectral data were acquired by using the hymap imaging system consisting of sensors that collect reflected solar radiation in 126 bands covering the 440–2500 nm wavelength range (bedini & tukiainen 2008). the spatial resolution was 4 × 4 m (tukiainen 2001). eight sites underlain by caledonian or post-caledonian rocks with known mineral occurrences (fig. 1) were tested. the project was financially supported by the greenland bureau of minerals and petroleum and the data were analysed by geus. here we provide a summary of the results. field work 2005–2009 ground checks were undertaken in 2005, 2008 and 2009 by geus (thomassen & tukiainen 2008). the field work by the authors was carried out from light-weight camps in co-operation with international molybdenum ltd. that explored the malmbjerg molybdenum deposit, and with other geus activities. the aim was to investigate hyperspectral anomalies and known mineral occurrences. the field work comprised ground measurements of rocks, minerals and their weathering products with a portable spectro-radiometer in order to determine their spectral character and to compare this information with the airborne data. in 2005 and 2008, a pima ii portable instrument was borrowed from other institutions, but for the 2009 season geus purchased an advanced spectro-radiometer, model fieldspec 3 hires. our investigations showed that there is good correlation between the airborne spectra and the field spectra, thus confirming the quality and stability of the airborne hyperspectral data. in general, sulphide minerals have poor to weak spectral response in the visible and near-infrared (vnir) and shortwave infrared (swir) spectral regions whereas their alteration products, such as malachite, cerussite, smithsonite and jarosite, are distinctly swir-active. however, apart from jarosite, these minerals are virtually non-existent in the region. in contrast, it appears from our study that the hyperspectral detection of typical hostand wall-rock alteration minerals (jarosite, white micas, phengite, kaolinite, dolomite etc.) provides an effective method to outline potential exploration targets. main results below we present some results that are of relevance for mineral exploration in the region. the reader is referred to harpøth et al. (1986) for a description of the mineral occurrences of the region and to henriksen et al. (2009) for a description of the regional geology. the locations of the described areas are shown in fig. 1. area 1 – wegener halvø. this horst-like peninsula exposes a complex pattern of fault blocks involving neoproterozoic to triassic sedimentary rocks, with stratabound base-metal mineralisation occurring in the permo-triassic section. disseminated mineralisation hosted by triassic sandstone and shale could not be detected by the airborne hyperspectral survey due to lack of alteration minerals. however, a close association of dolomitisation with upper permian, carbonate-hosted base-metal mineralisation is confirmed by our investigations, making the dolomite map a valuable exploration tool (fig. 2). area 2 – werner bjerge. a well-known porphyry molybdenum deposit is hosted in the malmbjerg granite stock, which is a unit of the palaeogene werner bjerge alkaline intrusive complex. the alteration zones surrounding the deposit constitute an important target for our investigation. conspicuous high-temperature, potassic and siliceous hydrothermal alteration outlined by muscovite and phengite is well displayed in the hyperspectral data. the high-temperature alteration apparently culminated in greisenisation of the malmbjerg granite stock, as exemplified by topaz-muscoviteenriched rocks in the stock roof and in the roofing permocarboniferous sandstone (fig. 3). in addition, the hyperspectral mapping of typical alteration minerals outlines a large number of potential exploration targets with low-temperature argillitic and propylitic alteration elsewhere in the alkaline intrusive complex. the hyperspectral mapping also defined a locality c. 1 × 1.5 km in size immediately south of werner bjerge, which displays many spectral similarities to malmbjerg. the accessible part of this anomaly at 1300 m was found to host a significant number of pyriteand fluorite-bearing trachytic application of airborne hyperspectral data to mineral exploration in north-east greenland tapani tukiainen and bjørn thomassen © geus, 2010. geological survey of denmark and greenland bulletin 20, 71–74. open access: www.geus.dk/publications/bull 7272 dykes and sheets a few metres thick in permo-carboniferous sandstone. the intrusive trachytic rocks host abundant wallrock fragments and are highly enriched in potassium and show elevated concentrations of tungsten (max. 35 ppm), molybdenum (max. 21 ppm) and thorium (max. 226 ppm). these rocks may represent the top of a porphyry system with an unexposed granite at a lower level. area 3 – mestersvig. the permo-carboniferous sandstones of this area host epithermal lead–zinc-bearing quartz veins, including the mined-out blyklippen deposit. the min100 km werner bjergewerner bjerge blyklippenblyklippen tnathorst land trailltraill øø j a m e s o n l a n d constable pynt gåseland malmbjergmalmbjerg milne land hinks land charcot land hudsonhudson landland claveringclavering ø ø wollastonwollaston forlandforland t canning land wegener halvø devondal kap simpson mestersvig werner bjerge p d m f malmbjerg blyklippen daneborg ’lauge koch’s gold mine’ parkinson bjerg traill ø ymer ø hold with hope clavering ø wollaston forland hudson land kong oscar f jord kejser franz joseph fjord inland ice greenland post-caledonian caledonian orogen ice quaternary, undifferentiated intrusive complexes palaeogene volcanic province sedimentary basins allochthonous thrust sheets structures boundary/unconformity thrust post-devonian main fault normal fault base, abandoned mine area with airborne hyperspectral data extensional fault and detachment separating thrust units basaltic sills and dykes basaltic plateau lavas jurassic–palaeogene triassic upper permian permo-carboniferous devonian granites (sensu lato) early neoproterozoic metamorphosed rocks mesoproterozoic metasediments palaeoproterozoic crystalline basement complexes archaean crystalline basement complexes neoproterozoic – lower palaeozoic sediments 4 pdmf 72°n 73°n 74°n 25°w 20° 1 4 2 3 5 6 7 8 fig. 1. geological map of north-east greenland showing areas covered by airborne hyperspectral data. the numbers refer to the areas discussed in the text. modified from henriksen & higgins (2008). 73 eralisation is accompanied by silicification and kaolinisation of the wall rocks, but this association did not clearly define the veins in the airborne survey. however, the airborne data reveal a distinct, c. 500 m wide zone of pervasive kaolinisation of the arkosic sandstone some 3 km north-east of the old mine. this could be related to unknown base-metal mineralisation of the blyklippen type. area 4 – kap simpson. the palaeogene kap simpson alkaline intrusive complex hosts a caldera structure displaying widespread and intensive hydrothermal alteration. pyrite is common and traces of molybdenite are known, with basemetal and niobium-bearing quartz and calcite veins in the host mesozoic sediments. the hyperspectral survey distinguished between low-temperature, fumarole-related alteration (montmorillonite-illite-jarosite and iron oxides) and high-temperature alteration (muscovite-phengite) associated with subvolcanic intrusions. high-temperature potassic alteration and greisen-like spectral signatures on a granite in the north-western part of the intrusive complex define a new exploration target with potential for porphyry-type mineralisation. area 5 – ymer ø. ymer ø hosts a number of e–w-striking, epithermal tungsten, antimony, gold, base-metal-bearing quartz veins in neoproterozoic sediments. samples from the antimonite-bearing veins returned up to 23.4% sb and 4.7 ppm au. these veins give a weak hyperspectral response, due to the presence of low-temperature argillitic minerals (kaolinite, illite and alunite) in the quartzitic wall rock. however, known scheelite-antimonite-bearing veins are not detected by the hyperspectral survey. this is due to lack of distinct alteration minerals other than quartz in the carbonate wall rock. distinct linear anomalies detected in the airborne data turned out to originate from 3–5 m thick, e–w-striking, unmineralised rhyolitic veins with kaolinite alteration or weathering products, probably related to the mineralising system. area 6 – hudson land. central hudson land exhibits various types of mineralisation in a complex pattern of proterozoic to palaeogene rocks transsected by a regional, n–strending structure, called the post-devonian main fault. quartz veins 0.1–1.0 m thick with greisen mineralisation returned up to 1.4% sn and 0.5% cu but were not depicted in the airborne data due to their modest size. in contrast, extensive, low-temperature, hydrothermal alteration with associated epithermal baseand noble-metal-bearing veins along the post-devonian main fault is clearly seen in the hyperspectral data. special attention was paid to a devonian granite stock at parkinson bjerg, which is surrounded by geochemical sn-w-mo-nb anomalies. the granite was found to have a pegmatitic core rich in quartz, fluorite and tourmaline, corresponding to a distinct tourmaline and phengite response in the airborne data. this granite is a potential source for the geochemical anomalies. rock samples returned up to 0.4% sn, 0.17% be and 0.1% y. area 7 – clavering ø. extensive rust zones caused by hydrothermal alteration and epithermal base-metal mineralisation along faults occur in proterozoic metasediments on clavering ø. the mineralisation was investigated in the 1930s when a test adit was excavated at the so-called ‘lauge koch’s gold mine’, which transpired to host a pyrite vein without anomalous gold. the mineralised structures are distinct in the hyperspectral mapping as lineaments with lowtemperature alteration minerals such as illite and jarosite. samples of massive, brecciated pyrite ‘ore’ with minor fluorite and galena returned <2 ppb au. area 8 – wollaston forland. prior to 2000, no mineralisation had been reported from wollaston forland, and the area was included in the airborne survey for biological reasons. the survey showed hyperspectral anomalies caused by jarositeand muscovite/sericite-rich zones or lithologies in the north-western part of the area, which is underlain by proterozoic metasedimentary rocks. a scree sample of pyritiferous paragneiss below an anomaly returned 4.1 ppm ir while another loose block returned 0.38 ppm au. surrounding blocks of ultramafic rocks indicate a magmatic component in the area with potential for platinum group mineralisation. cu-pb-(zn)cu-pb-(zn) cu-pbcu-pb cu-pbcu-pb cu-pbcu-pb cu-pbcu-pbcu-pb-(zn) cu-pb cu-pb cu-pb cu-pb fig. 2. perspective view from the south-east of the north slope of devondal showing dolomitic alteration (red) of upper permian limestone. the scree aprons enhance the surface impression of dolomite. known occurrences of cu-pb-(zn) mineralisation are indicated. background image is a colour composite of hymap bands 27(r), 18(g) and 4(b). no vertical exaggeration, relief is 600 m. 7474 concluding remarks the present study demonstrates that the hyperspectral method is well suited for mineral exploration in remote and mountainous arctic regions. the most obvious target for future use of this method in greenland seems to be the palaeogene igneous province that stretches for 1100 km along the east coast. this province should be investigated for hostand wall-rock alteration indicative of subvolcanic porphyry-type molybdenum mineralisation. references bedini, e. & tukiainen, t. 2008: using spectral mixture analysis of hyperspectral remote sensing data to map lithology of the sarfartoq carbonatite complex, southern west greenland. geological survey of denmark and greenland bulletin 17, 69–72. harpøth, o., pedersen, j.l., schønwandt, h.k. & thomassen, b. 1986: the mineral occurrences of central east greenland. meddelelser om grønland, geoscience 17, 139 pp. henriksen, n. & higgins, a.k. 2008: caledonian orogen of east greenland 70°n–82°n: geological map at 1:1 000 000 – concepts and principles of compilation. in: higgins, a.k., gilotti, j.a. & smith, m.p. (eds): the greenland caledonides: evolution of the northeast margin of laurentia. geological society of america memoir 202, 345–368. henriksen, n., higgins, a.k., kalsbeek, f. & pulvertaft, t.c.r. 2009: greenland from archaean to quaternary. geological survey of denmark and greenland bulletin 18, 126 pp. thomassen, b. & tukiainen, t. 2008: ground check of airborne hyperspectral anomalies in the greater mesters vig area, central east greenland. danmarks og grønlands geologiske undersøgelse rapport 2008/14, 85 pp. tukiainen, t. 2001: projects mineo and hypergreen: airborne hyperspectral data acquisition in east greenland for environmental monitoring and mineral exploration. geology of greenland survey bulletin 189, 122–126. malmbjerg granite stock b a fig. 3. perspective view from the south-west of the malmbjerg granite stock, no vertical exaggeration, relief is 500 m. the digital terrain model is based on lidar data from international molybdenum ltd. (re-sampled at 1 × 1 m resolution). a: minimum noise fraction transformed shortwave infrared data draped on the detailed digital terrain model. note the compositional zoning of the granite stock and intensive high-temperature alteration of the roofing rocks (hues of yellow and orange). b: orthoscopic lidar image draped on the detailed digital terrain model. the pixels mapped as topaz/tourmaline-bearing greisen are shown in red. the boundaries of the granite stock are indicated. authors’ addresses t.t., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: tt@geus.dk b.t., present address: avannaa resources ltd., dronningens tværgade 48, dk-1302 copenhagen k, denmark. geological survey of denmark and greenland. bulletin 10, 57-60 57 earthquake seismology is a rapidly evolving field that has provided a wealth of new information about deep geological structures on a regional scale over the last decade as well as information about dynamic processes in the earth. a major leap forward was the development of portable digital broad band (bb) seismographs around 1990. without any changes in configuration, these are able to record the signals from large distant earthquakes, as well as the signals from weak local events. bb seismographs typically cover a frequency range from 0.0083 hz to 50 hz, making them useful for studies ranging from the high frequency signals from explosions to the very low frequency oscillations following major earthquakes. the first seismological observatory in greenland was established in 1907 in qeqertarsuaq (gdh) and was in service for about five years (hjelme 1996). later, seismographs were established in ivittut (1927) and illoqqortoormiut (1928; sco), and the network has been regularly upgraded and expanded ever since (fig. 1). prior to the development of bb seismographs, each station was equipped with a set of seismographs with different frequency sensitivities in an attempt to cover both distant and local earthquakes. now just one small instrument is needed at each location. the geological survey of denmark and greenland (geus) operates four permanent bb seismographs in greenland (fig. 1), two of them in collaboration with foreign institutions. in addition to the permanent network, there are currently 13 temporary bb seismographs active in greenland, of which eight are operated by geus. three of the temporary seismograph stations were established as part of the danish continental shelf project (marcussen et al. 2004), and the remainder in connection with research projects. three temporary seismographs were deployed during 2005 as part of a research project aiming to resolve very deep regional structures in north greenland: the citronen fjord station (cfj, continental shelf project), and the stations in kullorsuaq (kul) and daneborg (dbg). the seismological service in greenland the study of earthquakes is international by nature. large earthquakes can be recorded world wide, and a good coverage of high quality data is crucial for an accurate determination of an epicentre. epicentral determination for large earthquakes is carried out at international data centres through collaboration of seismic services in almost all countries of the world. the primary task for the seismology group at geus is to maintain the seismological service for denmark and greenland (see gregersen et al. 2004). the seismological service runs the permanent network of seismographs, that collects high quality continuous data, analyses the data, and reports registered earthquake signals (phases) for local earthquakes as well as for regional and teleseismic events. the © geus, 2006. geological survey of denmark and greenland bulletin 10, 57–60. available at: www.geus.dk/publications/bull earthquake seismology in greenland – improved data with multiple applications tine b. larsen,trine dahl-jensen, peter voss,thomas møller jørgensen, søren gregersen and hans peter rasmussen n dag sco sfj ale nrs upn nuk attu sisimiut ivittut sum ffb cfj kia ass kulang kangerlussuaq is3 kag gdh umm asi ill ngr paa dy2 dbg is1 is2 soe hjo tule ilg nor greenland summer station active summer 2000 station deployed in 2005 long-term station long-term station currently not active permanent station abandoned analog site 250 km 60° 70° 80° 50° 40° 30° 90° 50°60°70°80° 40° 30° 20° 10° 0° 10° fig. 1. seismograph stations in greenland. 58 continuous digital waveform data are freely available directly from a geus server, as well as through international data centres. due to greenland’s size and geographical location, data from greenland are particularly important to the international seismological community. currently, the kangerlussuaq (sfj) seismograph and the seismograph at the summit ice camp (sum) are the only two in greenland with real-time data transmission. data from other stations are available with varying time delay. the phase readings are reported to international data centres such as the united states geological survey (usgs) and the international seismological centre (isc) through the weekly bulletin and the revised monthly bulletin. in 2005, geus reported a total of 3999 earthquake signals from the permanent seismograph network in greenland, including many of the aftershocks that continued for several months after the sumatra earthquake on 26 december, 2004, as well as the devastating 7.6 richter scale kashmir earthquake on 8 october, 2005 (fig. 2). the seismograph in kangerlussuaq was moved 300 m in february 2005 to a vault protected from local radar antenna disturbances. this was carried out in order to serve better the needs of the comprehensive test ban treaty organisation (ctbto). ctbto financed the move, and provides in addition the real-time satellite transmission of the continuous data. the instrumentation was supplemented in 2005 with an extra bb seismograph. the ctbto is an organisation under the united nations, and its goal is to detect nuclear explosions of more than 5 kg fissionable material anywhere on earth. an important tool for the organisation is seismological surveillance carried out by a primary and a secondary global network of designated secure seismograph stations. the kangerlussuaq seismograph is part of the secondary network, and was officially certified by ctbto on 24 november 2005. geus is denmark’s national authority for the ctbto. several minor earthquakes occurred in greenland in 2005. geus received reports that earthquakes were felt in qeqertarsuaq (gdh) on 30 march, in sisimiut on 23 july, at station nord (nor) on 30 august, and in attu on 23 october. tasiilaq (ang) is normally the place in greenland where most earthquakes are felt; however, no earthquakes were reported in 2005. including the reports of felt earthquakes noted above, geus has so far registered 20 earthquakes in greenland in 2005 (fig. 3). this number may increase when the preliminary earthquake catalogue is quality controlled and revised, and data from the temporary seismographs are recovered and included. studies on deep crustal and mantle structures denmark ratified the united nations convention on the law of the sea (unclos) in november 2004. after ratification of unclos, a country has ten years to collect the appropriate information and submit a claim for an extended continental shelf beyond 200 nautical miles. one of the potential claim areas is the continental shelf north of greenland (marcussen et al. 2004). three bb seismographs have been placed along the north greenland coast (fig. 1) with the aim of learning more about the thickness and structure of the crust of north greenland through receiver function analysis. this method has previously been used at many sites in greenland during the glatis (greenland lithosphere analysed teleseismically on the ice sheet) project, but all bb seismographs used in previous studies were located further south (dahl-jensen et al. 2003). the stations at frankfield bugt (ffb) and aftenstjernesø (ass) have been in operation since 2004, whereas the citronen fjord station (cfj) started recording data in 2005. these stations record signals from distant (teleseismic) earthquakes that can be processed to obtain information about crustal structure, e.g. depth to moho and even deeper structures. the data retrieved so far are of excellent quality. normally it is necessary to record data for more than a year in order to obtain a reliable estimate for the depth to moho, but for the aftenstjernesø station a depth to moho of 41 km and a vp/vs of 1.71 was determined using just a few months of data. the relatively sco hh z sco hh n minutes 0 10 20 30 40 50 60 sco hh e fig. 2. seismograms for the richter scale 7.6 kashmir earthquake on 8 october, 2005 recorded at illoqqortoormiut. the top seismogram shows the vertical motion, the middle seismogram motion in the north–south direction, and the bottom seismogram motion in the east–west direction. 59 large depth to moho at aftenstjernesø indicates that the station is on the rim of the precambrian shield area (fig. 3), and data from other locations are therefore necessary to resolve the question of crustal thickness in the franklinian basin that extends along the north greenland coastal region. surface waves from teleseismic earthquakes can provide information on deep geological structures between two seismographs recording the same earthquake. provided that two seismographs are located on the same great circle as the epicentre, differences in the signal recorded on the two instruments will reflect the geology affecting the wave propagating between the two stations. a dense web of epicentres and station pairs will make it possible to construct a seismic velocity model from depths of about 60 km to about 300 km. this kind of analysis was successfully carried out during the glatis project (darbyshire et al. 2004). similar studies were made previously by gregersen (1970, 1982) for the crust, using various kinds of seismic waves. in addition to the studies of velocity structure in the mantle, measurements of amplitude attenuation have been carried out (jørgensen 2005), resulting in preliminary maps of differences in the wave absorption properties of the deep structures. in order to take full advantage of the seismographs along the north coast of greenland and improve coverage inland north of 72°n, it was necessary to deploy two extra seismographs, one in the settlement of kullorsuaq (kul), and another at daneborg (dbg) that was previously used by the glatis project (fig. 1). both seismographs were installed in 2005 and will remain in operation until 2007. it is expected that a first-order model of the very deep velocity structures inland in north greenland can be constructed within the framework of the project. in 2006 another deep-structure seismological project will deploy a short profile of five bb seismographs across the safartoq kimberlite region south-east of sisimiut. the aim of this project is to reveal the crustal thickness in the area. the safartoq project will be initiated as a pilot study of the area in order to facilitate a proposal for a large-scale deep structure study with international partners. a larger array of instruments is necessary for studies of the deep lithospheric structures. glacial earthquakes glacial earthquakes are a peculiar type of seismological event primarily located in greenland. they were first described by ekström et al. (2003). the signals are dispersive, lack the characteristic pand swaves known from ordinary earthquakes and all signals appear to have their source beneath large glaciers. the focal mechanisms that can be calculated for the earthquakes are consistent with a large mass of ice moving abruptly downhill over an elastic medium (landslide model). in 2005 geocenter copenhagen funded a joint research project on glacial earthquakes involving geus and the institute of geography, university of copenhagen. through the glatis project geus possesses a unique seismological data set for greenland, previously used only for structural studies. when the majority of bb seismographs in greenland were installed, the existence of glacial earthquakes had not been recognised. however, after the start of the glacial earthquake project in 2005, it has become clear that the bb seismological data from greenland are a unique data set for the study of glacial earthquakes. 500 km north atlantic ocean permanent ice upper cretaceous – palaeogene sediments and basalts devonian–palaeogene sediments caledonian orogenic belt proterozoic sediments and volcanic rocks palaeoproterozoic orogenic belts archaean craton offshore basins with substantial sediment thicknesses major faults and thrusts lower palaeozoic sediments, franklinian basin n 5.0 4.0 3.0 2.0 fig. 3. map of earthquake locations in and around greenland for 19702005. only earthquakes with a magnitude of 2.0 or more on the richter scale are shown. the map is compiled from the geus earthquake database, supplemented with epicentre locations from the geological survey of canada, the united states geological survey and norsar in norway. glacial earthquakes have not previously been studied extensively using a local data set (fig. 4). one purpose of the project is to significantly improve the location accuracy for the earthquakes, and to improve the detection threshold by an order of magnitude from 4.7 on the richter scale to 3.7. detecting smaller earthquakes opens the possibility of revealing unknown ice streams within the greenland ice sheet. the occurrence of glacial earthquakes shows strong seasonal variation, with most earthquakes occurring during summer. this implies that the glacial earthquakes could contribute information as to how variations in climate parameters affect ice sheet dynamics. gps fieldwork is planned for summer 2006, when glacial earthquakes will be recorded simultaneously on both seismographs and differential gps. the objective of the project is to shed light on the mechanisms controlling the glacial earthquakes, with the possibility that glacial earthquakes can be used as a surveillance tool to study greenland ice sheet dynamics. acknowledgements geoforschungszentrum-potsdam (gfz), germany provides instrumentation and technical support to the seismograph in danmarkshavn, and together with the incorporated research institutions for seismology (iris), usa to the seismograph in kangerlussuaq. the bureau of minerals and petroleum, government of greenland, provided financial support to several of the projects described in this paper. geocenter copenhagen provides financial support for the project on glacial earthquakes. references dahl-jensen, t., larsen, t.b., woelbern, i., bach, t., hanka, w., kind, r., gregersen, s., mosegaard, k., voss, p. & gudmundsson, o. 2003: depth to moho in greenland: receiver-function analysis suggests two proterozoic blocks in greenland. earth and planetary science letters 205, 379–393. darbyshire, f.a., larsen, t.b., mosegaard, k., dahl-jensen, t., gudmundsson, o., bach, t., gregersen, s., pedersen, h.a. & hanka, w. 2004: a first detailed look at the greenland lithosphere and upper mantle, using rayleigh wave tomography. geophysical journal international 158, 267–287. ekström, g., nettles, m. & abers, g.a. 2003: glacial earthquakes, science 302, 622–624. gregersen, s. 1970: surface wave dispersion and crust structure in greenland. geophysical journal of the royal astronomical society 22, 22–39. gregersen, s. 1982: seismicity and observations of lg wave attenuation in greenland. tectonophysics 89, 77–93. gregersen, s., glendrup, m., larsen, t.b., voss, p. & rasmussen, h.p. 2004: seismology: neotectonics and structure of the baltic shield. bulletin of the geological survey of denmark and greenland 7, 25–28. hjelme, j. 1996: history of seismological stations in denmark and greenland. in: wahlström, r. (ed.) seismograph recording in sweden, norway – with arctic regions, denmark – with greenland, and finland. proceedings from the uppsala wiechert jubilee seminar, 49–57. uppsala: seismological department, uppsala university, sweden. jørgensen, t.m., 2005: attenuation of rayleigh waves in greenland, 78 pp. unpublished m.sc. thesis, university of copenhagen, denmark. marcussen, c., christensen, f.g., dahl-jensen, t., heinesen, m., lomholt, s., møller, j.j. & sørensen, k 2004: exploring for extended continental shelf claims off greenland and the faroe islands. bulletin of the geological survey of denmark and greenland 4, 61–64. authors’ address geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: tbl@geus.dk 60 time ale ang borg dag frb gdh hjo kbs kono nuk soe sum stack ang is3 soe hjo kul a b fig. 4. a: seismogram envelopes for a glacial earthquake, that occurred on 20 october, 2000. signal from the earthquake is shown in blue. the bottom seismogram is a stack of the twelve seismograms, relative to a chosen test epicentre. when the test epicentre is close to the true epicentre, stacking will produce a strong signal. for location of greenland stations see fig. 1. ale: alert, arctic canada. borg: borganes, iceland. frb: frobisher bay, arctic canada. kbs: kingsbay, svalbard; kono: kongsberg, norway. b: correlation for the signal in the stack, with the best correlation in blue. many test epicentres are modelled before the true epicentre can be determined, in this case centred in kangerlussuaq in east greenland (see fig. 1). geological survey of denmark and greenland bulletin 7, 2004, p 65-68 65 in 2004 the geological survey of denmark and greenland (geus) initiated a study of the origin and tectono-metamorphic evolution of greenstone belts and important regional structures in the central godthåbsfjord region, southern west greenland (fig. 1; hollis et al. 2004). like other archaean belts worldwide, these greenstone belts are locally host to gold mineralisation. their complexity requires a combination of detailed geological mapping, geochemistry, petrographic work and geochronological studies to develop models of their geological setting, evolution and gold mineralisation. regional geology the greenstone belts in the godthåbsfjord region represent remnants of archaean intrusive and extrusive mafic rocks, and minor chemical and clastic sediments. these were once thought to be dismembered parts of the same greenstone belt, but are now known to belong to unrelated groups of different ages, occurring within several distinct continental crustal terranes – the færingehavn, akia and tre brødre terranes (fig. 1; e.g. schiøtte et al. 1988). these terranes were amalgamated at 2950–2700 ma (e.g. friend & nutman 2005). the 3850–3300 ma færingehavn terrane largely comprises early archaean tonalitic orthogneiss, granite, ultramafic and gabbroic rocks, together with rocks of inferred supracrustal origin, all of which are intruded by mafic ameralik dykes. the 3200–2975 ma akia terrane comprises orthogneiss and greenstones. the 2826–2750 ma tre brødre terrane, which is in tectonic contact with the færingehavn terrane in many parts of the central godthåbsfjord region, and also with the akia terrane (fig. 1), is dominated by late archaean (c. 2825 ma) orthogneiss, but also includes metasedimentary and mafic volcanic rocks and gabbro-anorthosite. the boundary greenstone belts in the central godthåbsfjord region, southern west greenland julie a. hollis, jeroen a.m. van gool, agnete steenfelt and adam a. garde geological survey of denmark and greenland bulletin 7, 65–68 (2005) © geus, 2005 ata ne q fau lt iv in ng ui t fa ul t 3 1 2 q us su k go dt hå bs fjo rd store malene storø sermitsiaq bjørneøen qussuk peninsula q kangersuneq st o rø s he ar z o ne kobbefjo rd fau lt kobbefjord akia (c. 3200–2975 ma) tre brødre (c. 2826–2750 ma) færingehavn (c. 3850–3300 ma) kapisillik (c. 3075–2960 ma) tasiusarsuaq (c. 3000–2800 ma) amphibolite metasedimentary rocks qorqût granite archaean terranes 64°15´n 64°30´n 64°45´n 51°30´w 51°00´w fig. 1. overview map of the central godthåbsfjord region – the focus of mapping and sampling in 2004 – and distribution of greenstone belts (dominantly comprised of amphibolite). 1, bjørneøen greenstone belt; 2, qussuk greenstone belts; 3, storø greenstone belt. insets show distribution of major tectonostratigraphic terranes, and regional location of the mapping area. q, qingaaq. between the tre brødre and akia terranes is the high-grade mylonitic ivinnguit fault (c. 2720–2710 ma; e.g. friend et al. 1987), which can be traced continuously for tens of kilometres (fig. 1). the emplacement of the qôrqut sheeted granite complex (c. 2530 ma; fig. 1) in central godthåbsfjord postdated regional terrane assembly. post-archaean events include emplacement of proterozoic dolerite dykes (2400 –2200 ma) and regional-scale faulting, e.g. the ne–swtrending kobbefjord and ataneq faults (fig. 1). greenstone belts: structure, lithology and mineralogy three distinct nne–ssw-trending greenstone belts can be recognised in the central godthåbsfjord region: the storø, bjørneøen and qussuk greenstone belts (fig. 1). based on cross-cutting relationships and existing geochronological data, these belts lie in the akia (qussuk and bjørneøen belts) and the tre brødre (storø belt) terranes (fig. 1). the bjørneøen greenstone belt the bjørneøen greenstone belt (> 2975 ma, akia terrane) is a continuous nne–ssw-trending belt which shows tight to isoclinal folding, large-scale, upright folding, nw-directed thrusting along ne–sw-striking high strain zones, and late brittle to semi-ductile reworking of ductile structures. the ivinnguit fault forms a tectonic base to part of the bjørneøen greenstone belt (fig. 1). the belt is dominated by mafic amphibolites, with lesser intermediate schist, metagabbro and ultramafic rocks, and minor garnetand tourmaline-rich rocks of inferred sedimentary and volcanic origin. the mafic amphibolites are variably calc-silicate banded with rare pillow structures preserved in areas of low strain (fig. 2). schists of intermediate composition also preserve angular clasts of inferred volcanic origin in areas of low strain. thin packages of layered metasedimentary and metavolcanic rocks are seen at store malene, sermitsiaq and south-east bjørneøen (fig. 1) and commonly contain tourmaline-rich layers; in some cases they define refolded folds. they have been interpreted as stratabound syngenetic deposits, closely associated with the tungsten mineral scheelite (cawso4; appel & garde 1987). laminated quartzofeldspathic rocks may represent tuffs. relatively minor units of garnet-bearing quartzite and garnet-biotite schist ± muscovite, zoisite, chlorite and tourmaline occur, along with anthophyllite-cordierite-bearing magnesian schists and cordierite-bearing quartzitic rocks. this belt reached only upper greenschist to lower amphibolite facies conditions. the qussuk greenstone belts the qussuk greenstone belts (> 2975 ma, akia terrane) may be contiguous with the bjørneøen belt. they occur as nne–ssw-trending, isoclinally folded and steeply dipping units of kilometre-thickness. rock types present include intermediate and mafic amphibolites, with lesser amounts of metagabbro, ultramafic rocks and minor garnetand tourmaline-rich rocks of inferred sedimentary and volcanic origin. the amphibolite units are dominated by rocks of intermediate composition. in areas of low strain these preserve decimetre-scale angular clasts inferred to be volcanic in origin (fig. 3). these fragmental rocks commonly occur in association with laminated quartzofeldspathic rocks that may have been tuffs. rare quartz-rich garnet and locally tourmaline66 fig. 2. deformed pillow lavas on south-east bjørneøen, looking east, showing eastward younging direction. fig. 3. angular volcanic clasts in amphibolite of intermediate composition, qussuk peninsula. bearing rocks may be volcanogenic-exhalative or hydrothermal in origin, similar to the tourmaline-rich rocks from the bjørneøen belt. the storø greenstone belt the storø greenstone belt is interpreted as part of the tre brødre terrane because of intrusive contacts with inferred late archaean orthogneiss at the structural base of the belt on storø. the outcrop of this belt on storø – the focus of gold exploration in the region – is largely controlled by the 200–300 m wide storø shear zone and associated kilometrescale folds (fig. 4). the storø shear zone runs parallel to the older ivinnguit fault along the north-west coast of storø, and has excised part of the belt along this section (fig. 1). outside the storø shear zone, the belt preserves tight to isoclinal folds reflecting earlier deformation. the storø belt is dominated by mafic amphibolites with rare pillow structures, banded amphibolites, mica schists, sillimaniteand mica-bearing quartzites, thin iron formations, and lenses of ultramafic rocks. a gabbro-anorthosite body at the structural base of the belt forms the core of two large doubly plunging antiforms in the hanging wall of the storø shear zone. north and south of storø the belt thins and metasedimentary rocks are rare. pelitic schists throughout the storø belt preserve amphibolite facies mineral assemblages, which pre-date the storø shear zone and include garnet, biotite, plagioclase, quartz ± sillimanite ± muscovite ± cordierite ± graphite ± staurolite. primary depositional environments and plate tectonic setting in the bjørneøen belt the common occurrence of tourmaline (indicating boron-rich compositions) and scheelite in stratabound settings, together with the spatial association with pillow structures in mafic amphibolite, are consistent with a submarine exhalative setting with associated alteration via hot circulating fluids (cf. appel & garde 1987). in the qussuk belts tourmaline-rich quartzitic rocks may be indicative of a similar setting. mafic amphibolites from the qussuk and bjørneøen belts preserve flat rare-earth element (ree) spectra consistent with an oceanic plateau environment. however, intermediate rocks from both belts show elevated light ree spectra, and some mafic amphibolites from the bjørneøen belt show mixed signatures, suggestive of a possible island-arc environment. this indicates more than one important magmatic source for the qussuk and bjørneøen belts. the gabbro-anorthosite body at the structural base of the storø greenstone belt may be indicative of an oceanic, riftrelated setting, an interpretation supported by rare pillow structures in mafic amphibolites and the occurrence of ultramafic lenses that may represent former dykes. however, the inferred tuffite origin of intermediate amphibolite and the occurrence of quartzite units indicate a significant continental source. these observations fit with an island-arc proximal to a continent, or a change in environment during formation of this belt. further geochemical analyses and geochronological constraints are required. 67 fig. 4. east-closing synform in the footwall of the storø shear zone, on the south-western slopes of qingaaq mountain on storø (see fig. 1 for location). metadolerite nw se anorthosite1190 m orthogneiss storø shear zone in greenstone belt sub-vertical dipmoderate nw dip steep se dip 68 mineral occurrences and implications for mineral exploration regional studies have identified a nne–ssw-trending zone through central godthåbsfjord with anomalous gold, and gold pathfinder elements (steenfelt et al. 2003). this zone is parallel to the trace of the ivinnguit fault. high gold values have been reported from mica schists and amphibolite on storø (nunaminerals a/s, licence number 2002/07; e.g. appel et al. 2000), where exploration and prospecting activities continue. trace element data for rock, stream sediment and soil samples obtained in 2004 corroborate the high gold values previously found on bjørneøen (skyseth 1998; smith 1998), but also identify previously unrecognised gold occurrences on the qussuk peninsula. here values of 1.4 and 2.3 ppm au were measured in samples of respectively an intermediate amphibolite and a tourmaline-bearing quartzitic rock of inferred hydrothermal origin. this may be indicative of a link between early hydrothermal processes (possibly in a submarine volcanic environment) and gold mineralisation. mapping in 2004 and geochemical analyses of samples from altered zones related to shear zones and faults (including the ivinnguit and ataneq faults) did not identify a relationship between these structures and anomalous gold values. in fact, field relationships point to mineralisation early in the tectono-metamorphic history. geochemical results show gold occurrences in all three greenstone belts and – particularly in the case of the qussuk belt – not necessarily related to shear zone or fault structures. the nne–ssw-trending anomalous zone may simply be a reflection of the trend of greenstone belts in this region, rather than indicative of a relationship to nne–ssw-trending structures, such as the ivinnguit fault terrane boundary. further work further geochemical analyses of key major and trace elements are required to build on correlations and develop models for primary tectonic settings. important issues include the determination of absolute timing constraints on regional structures, the deposition history of sedimentary and volcanic rocks, and the metamorphic history. over 70 structurally constrained samples have been collected for this purpose. further fieldwork is also required to investigate some of the important findings, and to extend the study to greenstone belts further afield in the godthåbsfjord region. acknowledgements this project was financially supported by the bureau of minerals and petroleum, nuuk, greenland and undertaken in parallel with related geus projects. references appel, p.w.u. & garde, a.a. 1987: stratabound scheelite and stratiform tourmalinites in the archaean malene supracrustal rocks, southern west greenland. bulletin grønlands geologiske undersøgelse 156, 26 pp. appel, p.w.u., bliss, i.c., coller, d.w., grahl-madsen, l. & pedersen, j.s. 2000: recent au discoveries in archaean rocks of central west greenland. transactions of the institution of mining and metallurgy (sect. b: appl. earth sci.) 109, b34–b41. friend, c.r.l. & nutman, a.p. 2005: new pieces to the archaean terrane jigsaw puzzle in the nuuk region, southern west greenland: steps in transforming a simple insight into a complex regional tectonothermal model. journal of the geological society (london) 162, 147–162. friend, c.r.l., nutman, a.p. & mcgregor, v.r. 1987: late-archaean tectonics in the færingehavn – tre brødre area, south of buksefjorden, southern west greenland. journal of the geological society (london) 144, 369–376. hollis, j.a., van gool, j.a.m., steenfelt, a. & garde, a.a. 2004: greenstone belts in the central godthåbsfjord region, southern west greenland: preliminary results from field work in 2004. danmarks og grønlands geologiske undersøgelse rapport 2004/110, 110 pp. + 1 dvd. schiøtte, l., compston, w. & bridgwater, d. 1988: late archaean ages for the deposition of clastic sediments belonging to the malene supracrustals, southern west greenland: evidence from an ion probe u-pb zircon study. earth and planetary science letters 87, 45–58. skyseth, t. 1998: gold exploration on storø 1997, south west greenland, 25 pp. unpublished report, nunaoil a/s, nuuk, greenland (in archives of geological survey of denmark and greenland, geus report file 21601). smith, g.m. 1998: report on the structure and geometry of the gold mineralization at qingaq storø, nuukfjord, south west greenland, 13 pp. unpublished report, nunaoil a/s, nuuk, greenland (in archives of geological survey of denmark and greenland, geus report file 21602). steenfelt, a., moberg, e. & appel, p.w.u. 2003: geochemical data indicative of mineral occurrences. in: appel, p.w.u. et al.: economic potential of the greenstone belts in the nuuk area. general geology and evaluation of compiled geophysical, geochemical and ore geological data. danmarks og grønlands geologiske undersøgelse rapport 2003/94, 58–66. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: jho@geus.dk geological survey of denmark and greenland bulletin 12, 1-16 1 geological survey of denmark and greenland bulletin 12 · 2007 lithostratigraphy of the palaeogene – lower neogene succession of the danish north sea poul schiøler, jan andsbjerg, ole r. clausen, gregers dam, karen dybkjær, lars hamberg, claus heilmann-clausen, erik p. johannessen, lars e. kristensen, iain prince and jan a. rasmussen geological survey of denmark and greenland danish ministry of the environment 2 geological survey of denmark and greenland bulletin 12 keywords lithostratigraphy, biostratigraphy, north sea basin, palaeogene, neogene. cover complex fabric created by multiple small-scale sand intrusions (light) into dark mudstones – such enigmatic fabrics are commonly associated with the sand-rich units of the rogaland group in the siri canyon area, offshore denmark. the illustrated section of core is about 10 cm across and is from the lower tyr member (lista formation) in the cecilie-1b well (2346.8 m). photograph: jakob lautrup. chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, geological institute, university of aarhus; minik rosing, geological museum, university of copenhagen; finn surlyk, department of geography and geology, university of copenhagen scientific editors of this volume: jon r. ineson and martin sønderholm editorial secretaries: jane holst and esben w. glendal referees: paul van veen (norway) and robert o’b. knox (uk) illustrations: stefan sølberg digital photographic work: benny m. schark graphic production: knud gr@phic consult, odense, denmark printers: schultz grafisk, albertslund, denmark manuscript received: 29 august 2005 final version approved: 8 september 2006 printed: 29 june 2007 issn 1604-8156 isbn 978-87-7871-196-0 geological survey of denmark and greenland bulletin the series geological survey of denmark and greenland bulletin replaces geology of denmark survey bulletin and geology of greenland survey bulletin. citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 12, 77 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk and geografforlaget a/s filosofgangen 24, 1., dk-5000 odense c, denmark phone: +45 63 44 16 83, fax: +45 63 44 16 97, e-mail: go@geografforlaget.dk or at www.geus.dk/publications/bull © de nationale geologiske undersøgelser for danmark og grønland (geus), 2007 for the full text of the geus copyright clause, please refer to www.geus.dk/publications/bull 3 contents abstract . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5 introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7 geological setting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9 previous work . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10 material and methods. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 12 offshore and onshore lithostratigraphic nomenclature. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16 chronostratigraphy and biostratigraphy. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16 paleocene. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 20 eocene . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 20 oligocene . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 21 miocene . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23 lithostratigraphy . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 24 rogaland group . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 24 våle formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25 bor member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 30 lista formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 32 vile member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 36 tyr member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 39 ve member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 40 idun member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 41 bue member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 44 rind member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 45 sele formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 46 kolga member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 50 fur formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 51 balder formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 54 stronsay group . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57 horda formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57 hefring member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 63 westray group . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 63 lark formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 63 dufa member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 69 freja member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 71 acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 73 references . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 73 4 5 abstract schiøler, p., andsbjerg, j., clausen, o.r., dam, g., dybkjær, k., hamberg, l., heilmann-clausen, c., johannessen, e.p., kristensen, l.e., prince, i. & rasmussen, j.a. 2007: lithostratigraphy of the palaeogene – lower neogene succession of the danish north sea. geological survey of denmark and greenland bulletin 12, 77 pp. + 5 plates. as a result of a lithological, sedimentological and biostratigraphic study of well sections from the danish sector of the north sea, including some recently drilled exploration wells on the ringkøbing– fyn high, the lithostratigraphic framework for the siliciclastic palaeogene to lower neogene sediments of the danish sector of the north sea is revised. the sediment package from the top of the chalk group to the base of the nordland group is subdivided into seven formations containing eleven new members. the existing våle, lista, sele, fur, balder, horda and lark formations of previously published lithostratigraphic schemes are adequate for a subdivision of the danish sector at formation level. bor is a new sandstone member of the våle formation. the lista formation is subdivided into three new mudstone members: vile, ve and bue, and three new sandstone members: tyr, idun and rind. kolga is a new sandstone member of the sele formation. hefring is a new sandstone member of the horda formation. freja and dufa are two new sandstone members of the lark formation. danish reference sections are established for the formations, and the descriptions of their lithology, biostratigraphy, age and palaeoenvironmental setting are updated. __________________________________________________________________________________________________________ authors’ addresses p.s.*, j.a., k.d. & l.e.k., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. * present address: gns science, 1 fairway drive, avalon, p.o. box 30368, lower hutt, new zealand. e-mail: p.schioler@gns.cri.nz o.r.c. & c.h.-c., department of earth sciences, university of aarhus, høegh-guldbergsgade 2, dk-8000 århus c, denmark. g.d. & l.h., dong energy, agern allé 24–26, dk-2970 hørsholm, denmark. i.p. & e.p.j., statoil norway, forusbeen 50, n-4035 stavanger, norway. j.a.r., geological museum, university of copenhagen, øster voldgade 5–7, dk-1350 copenhagen k, denmark. 6 fig. 50 fig. 56a fig. 49 fig. 56b fig. 61 fig. 58 saxo-1 wessel-1 tordenskjold-1 eg-1 diamant-1 bertel-1 mona-1 karl-1 w. lulu-3,-1 cleo-1 augusta-1 amalie-1 tabita-1 gulnare-1 gwen-2 iris-1 baron-2 nora-1 elin-1w-1 ravn-1 falk-1 u-1 e-8 tove-1 john-flanke-1 alma-1 emma-1 edna-1 roxanne-1 ugle-1 frida-1 l-1 francisca-1 cecilie-1 connie-1 elna-1 siri-2 siri-3 sofie-1 floki-1 sandra-1 nolde-1 nini-1 nini-2 nini-3 d-1 vanessa-1 ibenholt-1 ida-1 r-1 c-1 k-1 f-1 inez-1 s-1 siri-1 v-1 g-1 deep-adda-1 adda-2,-1bo-1 north-jens-1 lulu-1 lulu-2 sten-1 gert-1 kim-1 lone-1 57°00' 4°00' 6°00' 56°00' 25 km 250 km c offee soil fault north sea denmark n s uk g nl p sir i c an yo n mid north sea high central graben a b norwegian–danish basin c entral g raben 100 km north polish strait east shetland platform fennoscandian shield scottish high v ik in g g ra be n rhenish massif bohemian massif jylland sjælland storebælt moray firth mid north sea high ringkøbing–fynhigh fig. 1. location maps showing the position of wells used in the study (a) and major structural elements in the greater north sea area (b) mentioned in the text. on the well map (a) are indicated the locations of the seismic sections shown in figs 49, 50, 56, 58 and 61. grey shading on this map indicates the margins of the siri canyon; grey shading inside the canyon indicates an area of positive relief within the canyon. ggggg, germany; nnnnn, norway; nlnlnlnlnl, netherlands; ppppp, poland; sssss, sweden; u ku ku ku ku k, united kingdom. 7 introduction intense drilling activity following the discovery of the siri field in 1995 has resulted in an improved understanding of the siliciclastic palaeogene sediment package in the danish sector of the north sea (fig. 1). many of the new wells were drilled in the search for oil reservoirs in sandstone bodies of paleocene–eocene age. the existing lithostratigraphy was established on the basis of data from a generation of wells that were drilled with deeper stratigraphic targets, with little or no interest in the overlying palaeogene sedimentary succession. this means that this early scheme does not include palaeogene sandstone units in the danish sector. in order to improve the understanding of the distribution, morphology and age of the palaeogene sediments, in particular the economically important sandstone bodies, a detailed study of this succession in the danish sector has been carried out. the main aim was to update the lithostratigraphic framework of the succession on the basis of new data from recently drilled wells. all of the widespread palaeogene mudstone units in the north sea were established with norwegian or united kingdom (uk) type wells. in the present work, these units have been maintained unchanged or with only slight modifications. danish reference wells have been established for the units, however, and lithological descriptions have been expanded to cover the characteristics of these units in the danish sector. many of the sandstone bodies recently discovered in the danish sector have a limited spatial distribution and are derived from sources different from those of most of the contemporaneous sandstone bodies in the norwegian and uk sectors; furthermore, the danish sandstone bodies probably neither overlap nor are in contact with the norwegian/uk sandstones. these units have therefore been established as new in the danish sector, and have been assigned danish type and reference sections. the lithostratigraphy presented herein (fig. 2) has its base at the top of the early paleocene (danian) ekofisk formation (chalk group). the top of the study section is at the unconformity between the late eocene – mid-miocene westray group and the mid-miocene to recent nordland group. oil companies operating in the north sea have collected a substantial amount of lithostratigraphic data on the palaeogene successions and a detailed lithostratigraphy has been developed for the danish and norwegian sectors (see e.g. hamberg et al. 2005). a number of informal lithostratigraphic units have been introduced that have subsequently found their way into academia and geological survey organisations. it has been the aim of the present work formally to define these new units. this has been done maintaining their original (albeit informal) names whenever feasible. it has not been the aim of this work to provide a sequence stratigraphic model for the palaeogene sediments in the central and eastern north sea; for this the reader is referred to michelsen et al. (1992, 1995, 1998), mudge & bujak (1994, 1996a, b), neal et al. (1994) and danielsen et al. (1997). the present contribution does not attempt to review the petroleum-related aspects of the palaeogene succession. information about this may be found elsewhere, for example in the annual reports from the danish energy authority. preliminary results from the present work, including a revised lithostratigraphic scheme, were previously published in a brief review paper (schiøler et al. 2005). the present contribution formally describes the new stratigraphic units suggested in the review paper and further documents the palaeogene – lower neogene lithostratigraphy in the danish sector of the north sea. 8 danian selandian thanetian sparnacian ypresian lutetian ma m id dl e eo ce ne lo w er e oc en e u pp er p al eo ce ne c ha lk g ro up r og al an d g ro up st ro ns ay g ro up w es tr ay g ro up st ro ns ay g ro up lo w er p al eo ce ne priabonian rupelian chattian aquitanian burdigalian langhian serravallian 15 20 25 30 35 45 50 55 60 65 m id dl e m io ce ne lo w er m io ce ne lo w er n eo ge ne pa la eo ge ne pa la eo ge ne u pp er o lig oc en e lo w er o lig oc en e u pp er e oc en e m id dl e eo ce ne lillebælt clay fm horda fm røsnæs clay fm fur fm f ølst fm stolle klint clay østerrende clay holmehus fm ve mb bue mb rind mb æbelø fm danian limestone ekofisk fm våle fm li st a fo rm at io n vile mb sele fm balder fm bartonian lutetian 40 søvind marl fm søvind marl fm viborg fm linde clay branden clay vejle fjord fm hodde fm gram fm lark fm (undivided) horda fm se ri es sy st em st ag e d en m ar k on sh or e danish north sea kolga mb hefring mb tyr mb bor mb 49.0 55.5 54.5 57.9 60.0 41.3 37.0 33.7 28.5 23.8 20.5 16.4 14.8 dufa mb odderup fm bastrup sand ribe fm freja mb idun mb n or dl an d g ro up arnum fm kerteminde marl lellinge greensand 9 geological setting the danish sector of the north sea is situated in the central and eastern north sea and comprises three major structural elements: the central graben, the norwegian–danish basin (the eastern part of the northern north sea basin of rhys 1974) and the ringkøbing–fyn high (fig. 1; the geographic terminology and names of structural elements in the north sea used herein are adapted from rhys 1974, rønnevik et al. 1975, deegan & scull 1977 and fyfe et al. 2003). the western boundary of the danish sector largely coincides with the eastern boundary of the mid north sea high, the southern boundary largely coincides with the southern limit of the ringkøbing–fyn high, and the northern boundary is in the norwegian– danish basin. this basin as well as the ringkøbing–fyn high are early permian structures. active rifting occurred in the central graben from the middle to late jurassic along pre-established palaeozoic fault trends. major tectonic activity around the palaeozoic and jurassic structures had largely ceased by late cretaceous time, and the sediment basin below the central north sea was largely characterised by regional subsidence (ziegler 1981). during the late cretaceous to danian sea-level high, pelagic chalk sediments draped the structural highs and the northern and southern north sea basins became one north sea basin delimited by the fennoscandian shield to the north-east, the rheinish–bohemian massif to the south and the british massifs, highs and platforms to the west (see ziegler 1981 fig. 16 for details). chalk sedimentation continued through to the end of the danian stage when it gave way to hemipelagic and siliciclastic sedimentation. this was probably caused by uplift of the basin margins to the west and east (ahmadi et al. 2003). however, most of the siliciclastic sediments were derived from the scottish high and the east shetland platform, uplifted by the iceland plume (ahmadi et al. 2003). by the time of peak uplift, in the mid-thanetian, large sand systems were building out towards the central north sea. most sediment came from the west, but the siri canyon system, a depression in the top chalk surface, was fed from the fennoscandian shield in the north-east and north (fig 1; ahmadi et al. 2003; hamberg et al. 2005). thermal subsidence centered above the central graben continued through the eocene as sea level fell and the temperature decreased. shallow-marine sediments characterised the margins of the north sea basin, especially its western margin, whereas basinal mudstone continued to accumulate in the basin centre and in the eastern part of the basin (joy 1996). inversions controlled by compression between the atlantic spreading zone to the northwest and the orogenesis of the alps to the south added to further uplift of the basin margins and submarine fans and turbidites were deposited near the centre of the basin (jones et al. 2003). during the oligocene, the north sea basin became part of a larger nw european basin. connection with the north atlantic broadened and enhanced communication with the oceanic water mass to the north-west, whereas the connection to the south through the north polish strait became closed for the deep water (fyfe et al. 2003). glacio-eustatic sea-level changes became more frequent and controlled the sedimentary cycles. the eastward progradation direction of the paleocene and eocene sediments gave way to sediment supply from the european massifs to the far south (fyfe et al. 2003). continued subsidence above the mesozoic rift structures created accommodation space for thick sediment packages of basinal mudstones, and few sandstone units reached the basin depocentre above the mesozoic rifts (fyfe et al. 2003). in the neogene epoch, sediment started to be derived from the fennoscandian shield to the north, and the progradation direction changed to the south-west and west in the danish sector of the north sea. facing page: fig. 2. lithostratigraphic column for the palaeogene and lower neogene of the danish north sea sector showing the approximate correlation with danish onshore stratigraphic units. timescale from hardenbol et al. (1998), except for the age of the paleocene–eocene boundary, which is adapted from berggren & aubry (1996) and the age of the sparnacian–ypresian boundary, which is from aubry et al. (2003). stratigraphy and ages of pre-chatian onshore lithostratigraphic units are based on heilmann-clausen (1995) and clemmensen & thomsen (2005). post-rupelian onshore stratigraphy and ages are from dybkjær & rasmussen (2000) and rasmussen (2004a). f, fur formation. 10 previous work the permian to recent lithostratigraphy of the north sea was described in two pioneering stratigraphic works. rhys (1974) provided an overview of the structural elements of the north sea and gave a brief description of the palaeogene sediments. deegan & scull (1977) compiled a detailed lithostratigraphic subdivision and lithological description for the central and northern north sea (figs 3, 4). they subdivided the siliciclastic palaeogene, neogene and quaternary sediments into five major groups: the montrose, moray, rogaland, hordaland and nordland groups. the montrose and moray groups established for the outer moray firth – forties area are proximal equivalents to the rogaland group and are not present in the danish sector, whereas the rogaland, hordaland and nordland groups have widespread distribution in the danish sector. the succession of major mudstone formations contained within the three basinwide groups has formed the backbone of all subsequent lithostratigraphic schemes for the central and northern north sea, including that of the present contribution. the post-danian cainozoic succession of the danish central graben was divided into seven informal units by kristoffersen & bang (1982). the palaeogene comprised five units: north sea marl and cen-1–4 (fig. 4). the ranks of the units were not stated. although descriptions and interpretation of the cen units were detailed, they are essentially informal and have been little used. a revised lithostratigraphy for the palaeogene and neogene of the norwegian north sea sector was published by hardt et al. (1989). their lithostratigraphic scheme includes a number of new palaeogene and neogene sandstone bodies observed in the norwegian and british sectors of the north sea (fig. 4). some of the names of the new sandstone units established by hardt et al. (1989) were subsequently used informally for comparable sandstone units discovered in the danish sector. mudge & copestake (1992a, b) presented a revised palaeogene stratigraphy for the outer moray firth and northern north sea basins. in their papers they redefined the moray and montrose groups of deegan & scull (1977) and abandoned the rogaland group. the authors also demoted the previously established sandstone formations within the two former groups to the rank of members. besides, in an innovative approach they allowed for a greater influence of biostratigraphic data on the characterisation of the various lithostratigraphic units, an approach which is also followed herein. knox & holloway (1992) updated the lithostratigraphic scheme for the palaeogene in the british and norwegian central and northern north sea (figs 3, 4). the authors followed mudge & copestake (1992a, b) in abandoning the rogaland group of deegan & scull (1977), and used mudge & copestake’s revised definition of the montrose and moray groups for the central north sea as well. furthermore, the thick and hitherto undivided hordaland group was subdivided into two new groups, the stronsay group succeeded by the westray group, each containing a distal and a proximal formation. the two distal formations of the two groups, the horda and lark formations, together constitute the bulk of the palaeogene sediments in the danish sector of the north sea and are adopted herein (figs 2–4). although sandstone units occur in both the horda and lark formations in the danish sector, the two proximal sandstone formations of the fig. 3. correlation chart showing the approximate correlation between key lithostratigraphic schemes for the central and eastern north sea at group and formation levels. fu r deegan & scull (1977) hardt et al. (1989) knox & holloway (1992) this study chalk group chalk group chalk group hordaland group lista unnamed unit/ våle sele balder balder balder lista maureen montrose group moray group lark m ou sa sk ad e horda sele rogaland group rogaland group stronsay group westray group stronsay group westray group lista sele horda lark våle nordland group nordland group nordland group 11 stronsay and westray groups, the mousa and skade formations, are absent from the danish sector. following detailed analysis of new, high-resolution seismic surveys covering the succession in the eastern north sea area, efforts were focused on establishing a sequence stratigraphic subdivision of the palaeogene–neogene sediment package. the sedimentary succession was interpreted in a series of publications from a working group at the university of aarhus (e.g. michelsen et al. 1992, 1995, 1998; michelsen 1993; danielsen et al. 1997; huuse & clausen 2001). the result of that work was a subdivision of the palaeogene to mid-neogene sediment package covered by the present work into six genetic units (fig. 4). the sequence stratigraphy of the upper oligocene to miocene in the eastern north sea was dealt with by rasmussen (2004b). further sequence stratigraphic contributions covering the larger north sea basin including the british and norwegian sectors are given by armentrout et al. (1993), mudge & bujak (1994, 1996a, b) and neal et al. (1994). coastal onlap basinward 7 michelsen et al. (1998) 6 6.3 6.2 6.1 5.4 5.3 5.2 5.1 4.4 4.3 4.2 4.1 1.2 1.1 5 4 3 2 1 nordland group hordaland group balder sele lista unnamed unit ekofisk cen-5 cen-4 cen-3 cen-2 cen-1 north sea marl chalk-6 nordland group hordaland group balder sele sele lista lista våle ekofisk nordland group lark horda balder tay m o u sa s ka d e fr ej a k ol ga f u r r in d id un ty r bo r sele lista l is ta fo rt ie s c ro m ar ty m ey v ad e fo rt ie s a n d re w h ei m d al h er m o d f is ke b an k f is ke b an k f ri gg r og al an d g ro u p maureen m au re en ekofisk ekofisk ekofisk nordland group lark horda balder sele bue ve vile vålevåle deegan & scull (1977) northern north sea central north sea kristoffersen & bang (1982) hardt et al. (1989) knox & holloway (1992) this study nordland group balder ty grid ekofisk frigg h ef ri ng d uf a skade fig. 4. correlation chart showing approximate correlation between key lithostratigraphic schemes for the central and eastern north sea and the norwegian part of the northern north sea at formation and member levels. the sequence stratigraphic subdivision of michelsen et al. (1998) is added for comparison. sandstone-dominated units indicated in yellow. 12 material and methods the present lithostratigraphic subdivision represents the combined results from studies of petrophysical logs, biostratigraphy and seismic profiles, cuttings samples and cored sections. petrophysical logs from c. 70 wells in the danish sector have been scrutinised (see fig. 1 for well locations). the wells have been correlated using petrophysical logs, predominantly gamma-ray and sonic logs. five log panels form the basis for the log correlation (plates 1–5). lithostratigraphic well correlation has been supported by biostratigraphic data: biostratigraphic reports from 29 wells have been re-assessed with the aim of identifying key micropalaeontological and palynological events that occur consistently within the study area (taxa used are planktonic and benthic foraminifers, diatoms, radiolaria, sporomorphs and dinoflagellate cysts). moreover, biostratigraphic sample suites from 11 north sea wells have been prepared at the geological survey of denmark and greenland in order to further determine the biostratigraphic event succession. the bulk of material studied for biostratigraphy is based on cuttings samples, and only few table 1. well data for the new type and reference wells in the danish sector of the north sea augusta-1 cecilie-1 cleo-1 connie-1 e-8 f-1 floki-1 francisca-1 frida-1 inez-1 k-1 mona-1 nini-3 sandra-1 siri-1 siri-2 siri-3 tabita-1 bor mb(t), bue mb(t),ve mb(t) bor mb(r), tyr mb(r) bue mb(r), lista fm(r), ve mb(r), vile mb(r) idun mb(t), rind mb(t) bue mb(r), lista fm(r), ve mb(r), vile mb(r), våle fm(r) dufa mb(r) hefring mb(t) freja mb(t) freja mb(r) dufa mb(t), fur fm(r) fur fm(r) balder fm(r), horda fm(r), lark fm(r) kolga mb(r), tyr mb(t) rind mb(r) horda fm(r), lark fm(r), sele fm(r), våle fm(r) idun mb(r) balder fm(r), kolga mb(t), vile mb(t) sele fm(r) 56°17´57.40´́ n 04°24´04.64´́ e 56°24´23.73´́ n 04°45´42.00´́ e 56°23´23.54´́ n 04°25´22.70´́ e 56°24´28.34´́ n 04°42´30.36´́ e 55°38´13.42´́ n 04°59´11.96´́ e 57°01´53.4´́ n 06°54´28.6´́ e 56°27´48.58´́ n 05°16´47.11´́ e 56°22´27.95´́ n 04°48´05.30´́ e 56°17´14.15´́ n 05°01´50.20´́ e 56°50´28.39´́ n 06°57´41.62´́ e 57°07´37.74´́ n 07°09´43.11´́ e 56°16´35.94´́ n 04°00´15.81´́ e 56°41´31.96´́ n 05°24´12.35´́ e 56°35´13.33´́ n 05°01´35.19´́ e 56°29´11.10´́ n 04°54´57.49´́ e 56°29´40.53´́ n 04°52´13.26´́ e 56°30´34.92´́ n 05°03´48.27´́ e 56°13´37.50´́ n 04°23´47.56´́ e 04.03.2001 2991.0 mdrt 37.8 rt 65 15.10.2000 2361.0 mdrt 37.8 rt 59.4 06.02.1984 4866.1 mdkb 40.5 kb 63.1 02.02.2001 2351.8 mdrt 37.8 rt 61.5 08.04.1994 2527.4 mdkb 36.6 kb 43.6 06.10.1968 2421.6 mdkb 37.19 kb 40.8 29.08.2000 1878 mdrt 35.8 rt 53.2 20.07.1998 1888.5 mdrt 36.4 kb 60 26.07.1997 2274 mdrt 39.0 rt 54.3 11.09.1977 1983.9 mdkb 35.1 kb 35.4 22.01.1970 2292.4 mdkb 37.2 kb 56.4 03.10.1982 4241.6 mdkb 36.6 kb 65.5 12.01.2001 1851.2 mdrt 37.3 rt 58.2 18.06.1998 2139 mdrt 36 kb 65 28.11.1995 2220 mdkb 23 kb 60 03.08.1996 2297.5 mdrt 36.6 rt 60.6 30.08.1996 2171.5 mdrt 36.6 rt 60.1 10.09.1983 4353 mdkb 40 kb 65 dong e&p a/s dong e&p a/s chevron petroleum co. dong e&p a/s maersk oil & gas a/s gulf oil company kerr-mcgee int. aps dansk operatørselskab i/s dansk operatørselskab i/s chevron petroleum co. california oil co. chevron petroleum co. dong e&p a/s statoil e&p a/s statoil e&p a/s statoil e&p a/s statoil e&p a/s statoil e&p a/s type (t) or reference (r) well coordinates operator spud date td (logger’s kb/rt elevation water for listed units: depth in m) (m above msl) depth (m)well fm: formation. mb: member. mdrt: measured depth below rotary table. mdkb: measured depth below kelly bushing. 13 fig. 5. chronostratigraphy and biostratigraphy of the paleocene – middle miocene. a: paleocene–eocene. b: eocene–oligocene. c: oligocene – middle miocene. calibration of chronostratigraphic units follows hardenbol et al. (1998), berggren & aubry (1996) for the paleocene–eocene boundary and aubry et al. (2003) for the sparnacian–ypresian boundary. key dinoflagellate datums are calibrated mainly using age estimates from hardenbol et al. (1998) and williams et al. (2004). key microfossil datums are calibrated via their correlation with calibrated dinoflagellate datums as suggested by mudge & bujak (1996b), using age estimates from hardenbol et al. (1998) and williams et al. (2004). the combined event succession is correlated with the north sea microfossil zonation of king (1989) and lithostratigraphic units treated herein. in the microfossil event column, the planktonic foraminifer events appear in normal font, benthic foraminifers in italics; diatoms and radiolarians are underlined. senoniasphaera inornata palynodinium grallator, dinogymnium spp. alisocysta reticulata abundant p. pyrophorum isabelidinium? viborgense p. pyrophorum, p. australinum acme a. gippingensis alisocysta margarita common cerodinium wardenense apectodinium augustum apectodinium augustum, acme apectodinium spp. acme d. oebisfeldensis, influx inaperturopollenites spp., common h. tubiferum deflandrea oebisfeldensis dracodinium condylos nsp6 (pars) nsp5b nsp5a nsp4 nsp3 nsp2 nsp1 a b c nsb4 (pars) nsb3b nsb3a nsb2 nsb1 a b c horda balder sele ekofisk tor våle bue ve vile lista planktonic foraminifers benthic foraminifers diatoms and radiolaria planktonic microfossils north sea biozones (king 1989) lithostratigraphy selected biostratigraphic events used in the present studygeochronology ma a chronostratigraphy (berggren et al. 1995) c re ta ce ou s (p ar s) pa le oc en e eo ce ne (p ar s) u pp er (p ar s) lo w er u pp er lo w er (p ar s) fm mbdinoflagellate cysts ypresian (pars) maastrictian (pars) thanetian sparnacian selandian danian 54.5 55.5 57.9 60.0 65.0 50 55 60 65 uvigerina batjesi turrillina brevispira gaudryina hiltermanni common subbotina ex gr. linaperta fenestrella antiqua, foraminifers very rare impoverished benthic agglutinated assemblage common globoconusa daubjergensis globanomalina cf. compressa, s. trivialis increasing diversity of calcareous foraminifers reappearance of planktonic foraminifers increasing diversity of calcareous benthic foraminifers cenodiscus spp., cenosphaera spp. cretaceous foraminifers common f. antiqua and coscinodiscus morsianus pseudotextularia elegans benthic microfossilsstageseries 14 35 40 50 45 lark formation rupelian (pars) o lig oc en e (p ar s) eo ce ne ( pa rs ) priabonian 41.3 bartonian lutetian ypresian (pars) nsb7a nsb6b nsb6a nsb5c nsb5b nsb5a nsb4 nsb3a nsb2 (pars) nsb3b planulina costata pseudohastigerina spp. abundant radiolaria (cenosphaera spp.), cyclammina amplectens lenticulina gutticostata, spiroplectammina spectabilis balder horda eatonicysta ursulae diphyes ficusoides areosphaeridium michoudii heteraulacacysta porosa diphyes colligerum areosphaeridium diktyoplokum common e. ursulae phthanoperidinium clithridium globigerinatheka index cibicidoides truncanus vaginulinopsis decorata 49.0 37.0 33.7 uvigerina batjesi turrillina brevispira gaudryina hiltermanni common subbotina patagonica dracodinium condylos deflandrea oebisfeldensis acme d. oebisfeldensis, influx inaperturopollenites spp., common h. tubiferum fenestrella antiqua, foraminifers very rare cerebrocysta bartonensis uvigerina germanica karrulina conversa corrudinium incompositum sele (pars) nsp9b nsp9a nsp8c nsp8b nsp8a nsp7 nsp6 nsp5b nsp4 (pars) nsp5a lo w er ( pa rs ) u pp er m id dl e lo w er ( pa rs ) planktonic foraminifers benthic foraminifers diatoms and radiolaria planktonic microfossils north sea biozones (king 1989) lithostratigraphy selected biostratigraphic events used in the present study chronostratigraphy (berggren et al. 1995) fm mbdinoflagellate cysts benthic microfossilsstageseries geochronology ma b fig. 5b. chronostratigraphy and biostratigraphy of the eocene–oligocene. 15 lark nsp9a (pars) nsb6b (pars) uvigerina germanica karrulina conversa nsp14b nsb13a nsp14a nsb12c nsp13 nsb12b nsb12a nsp12 nsb11 nsp11 nsp10 nsb10 nsb9 nsp9c nsb8c nsb8b nsb8a nsb7b nsb7a nsp9b aulacodiscus allorgei turrillina alsatica bolboforma spiralis asterigerina staeschei, elphidium inflatum, meonis pompilioides uvigerina tenuipustulata plectofrondicularia seminuda aulacodiscus insignis quadrata (small), b. antiqua, g. girardana common elphidium subnodosum, common paragloborotalia nana rotaliatina bulimoides “turborotalia” ampliapertura common a. guerichi, paragloborotalia opima s.s. bolboforma metzmacheri pararotalia canui aulacodiscus insignis quadrata (large) spirosigmoilinella compressa cibicidoides mexicanus gyroidina mamillata wetzeliella gochtii phthanoperidinium amoenum chiropteridium spp. membranophoridium aspinatum distatodinium biffi cordosphaeridium cantharellus apteodinium spiridoides caligodinium amiculum thalassiphora pelagica hystrichokolpoma cinctum rhombodinium draco corrudinium incompositum achilleodinium biformoides enneadocysta pectiniformis burdigalian aquitanian m io ce ne (p ar s) lo w er chattian u pp er rupelian (pars) lo w er (p ar s)o lig oc en e (p ar s) m id dl e langhian serravallian 28.5 23.8 20.5 16.4 14.8 11.2 cousteaudinium aubryae nordland group bulimina elongata bolboforma clodiusi p. comatum 15 20 30 25 cannosphaeropsis passio tortonian (pars)u pp er (p ar s) planktonic foraminifers benthic foraminifers diatoms and radiolaria planktonic microfossils north sea biozones (king 1989) lithostratigraphy selected biostratigraphic events used in the present study chronostratigraphy (berggren et al. 1995) fm mbdinoflagellate cysts benthic microfossilsstageseries geochronology ma c fig 5c. chronostratigraphy and biostratigraphy of the oligocene – middle miocene. 16 core samples have been available. as the use of stratigraphic lowest occurrences (lo) of taxa in cuttings samples may be hampered due to downhole caving, the event succession comprises almost exclusively stratigraphic highest occurrences (ho) of taxa (a single significant lo is included in the succession). the event succession is shown in fig. 5a–c; its correlation with international and north sea biozones is shown in fig. 6a–c. seismic sections from the 2-d and 3-d seismic surveys cgd85, dk-1, rtd81–re94, ucg96 and ucge97 have been used to further support the well correlation and to map the stratigraphic units in areas with only scattered well coverage. the combined results from the correlation and mapping procedures are presented as isochore maps for individual stratigraphic units. inspection of cuttings samples from 16 key wells supplemented with sedimentological studies of cored intervals from 23 wells have formed the basis for the lithological and sedimentological descriptions of the units. the well depths mentioned in the lithostratigraphy section are loggers’ depths measured either from rotary table (mdrt) or kelly bushing (mdkb). supplementary data for new type and reference wells are provided in table 1. the names assigned to the new lithostratigraphic units defined herein are derived from nordic mythology and thus follow the nomenclatural tradition previously established for the norwegian north sea (isaksen & tonstad 1989). it should be noted that the micropalaeontology-based palaeoenvironmental terminology used herein was originally developed for a passive margin situation (e.g. the terms ‘neritic’ and ‘bathyal’ to indicate the physiographic zones ‘shelf ’ and ‘shelfslope’, respectively). its application herein to the epicontinental north sea basin solely relates to depositional depth. offshore and onshore lithostratigraphic nomenclature there is a high degree of lithological similarity between the palaeogene–neogene mudstone succession in danish offshore boreholes and that in onshore exposures and boreholes. however, the status of the danish onshore units is quite varied since many units were named before a standard for description of a lithostratigraphic unit was established; some fulfil these requirements, whereas others are still informal. if a previously established onshore unit and an offshore unit can be demonstrated to be identical (e.g. the holmehus formation and the new ve member proposed herein), the name of the onshore unit theoretically has priority over the name of the offshore unit (salvador 1994). in other cases, names of offshore units can be argued to have priority over onshore units (e.g. sele and balder formations over ølst formation). however, in order to acknowledge the traditional distinction between offshore and onshore stratigraphic nomenclature, the two sets of nomenclature are kept separate herein. whenever possible, comments are given in the text to explain the relationship between offshore and onshore danish stratigraphic nomenclature. a correlation between the two sets of nomenclature is shown in fig. 2. chronostratigraphy and biostratigraphy age assessment of the lithostratigraphic units in the north sea sedimentary succession is based on correlation between key biostratigraphic events encountered in the units and the calibrated standard chronostratigraphy published by berggren et al. (1995), with modification for the paleocene–eocene boundary following ratification of its position by the international union of geological scientists (aubry et al. 2002). the key events are from biostratigraphic zonation schemes established for the north sea area. planktonic and benthic microfossils are covered by the zonation schemes of king (1983, 1989; figs 5a–c, 6a–c). dinoflagellates from the paleocene and eocene epochs are covered by the zonation scheme of mudge & bujak (1996b; fig. 6a, b); the oligocene and miocene epochs are covered by the zonation schemes of costa & manum (1988) with modifications by köthe (1990, 2003; fig. 6b, c). key events from these schemes used in this study are listed in fig. 5a–c. for the dinoflagellate events, geochronological calibration has been largely established using age estimates from hardenbol et al. (1998), munsterman & brinkhuis (2004) and williams et al. (2004). for events not mentioned in these three publications, the works of mudge & bujak untitled 148 dynamic development of the thin-skinned thrust faulting the dynamic development of the thin-skinned thrust faulting in the rubjerg knude glaciotectonic complex is presented as a sequence of restoration stages. thus, the progressive deformation of thin-skinned thrust faulting and related syntectonic depositional developments are illustrated in sequentially restored crosssections beginning with the proximal moserende section and concluding with the ulstrup section in the most distal part of the thrust-fault complex. the basis of each restoration sequence is the balanced profile (plate 2a), and the end stage is identical with the thrustfault cross-section (plate 2b), including the interpretation of the unexposed ramps and flats in the subsurface. the most proximal sections, the mårup kirke and the ribjerg sections were interpreted individually in the preceding chapters, and are not included here. in a summary scheme (see fig. 123), it is concluded that the dynamic development was a process of continuous progressive deformation. thus, although the following description is concentrated on the individual sections, it should be kept in mind that there is overlap between sections, and that the whole system was mobile. thus a displacement of 5 m on one thrust might be followed by 10 m on a more proximal thrust and 7 m on a more distal thrust depending on the local conditions. this is the reason why a number of displacements appear to be out-of-sequence, but within limits that respect the lowest décollement level, and that displacements along the most distal, leading-edge thrusts were the last to be activated. it is therefore also evident that displacement along a leading-edge ramp may correspond to a translation along a flat in a proximal section. moserende section the thrust-fault development in the moserende section is regarded as normal progressive piggyback thrusting from the proximal towards the distal part. the mr12 thrust sheet was probably the first to be thrust onto the relatively thinner piggyback basin on the back of mr11 after a c. 10 m thickness of rubjerg knude formation sediments had been deposited. this is included in the first stage of the sequential restoration (fig. 112, stage 1). a total of eight stages have been dif ferentiated, of which stages 1–6 are illustrated in fig. 112. the stage preceding the deformation is shown in plate 2a, and the final stage terminating the deformation is reconstructed in plate 2b. moserende stage 1. the initial thrusting started with 40 m displacement of mr12 over the back of what was to become mr11. this thrusting was rooted down to the 20 m intermediate décollement level. during accumulation of a 20 m thick succession of sediments in the piggyback basins above the mr13 and mr12 sheets, the thrusting progressed with ramping of mr11 over the rubjerg knude formation on top of what was to become mr10. this thrusting involved ramping and translation of the lower segments of mr13– mr11 from the 30 m flat level onto the 20 m flat level. the trailing-end segments of the moserende section were contemporaneously over-thrust by mk01, the frontal thrust of the mårup kirke section, which is rooted in the 40 m décollement level. the accumulated displacement of thrusting of mr13, mr12 and mr11 is estimated at about 150 m. moserende stage 2. thrusting of mr09 initiated this stage. the mr09 thrusting ramped up from the 40 m décollement level, and a single duplex formed during stacking of the lower mr09 thrust segment. the mr09 sheet was displaced c. 40 m over the mr08 piggyback basin. contemporaneously, mr10 was thrusted over mr09 and the mr10 hanging-wall flat extended from the top flat level down to the 40 m décollement level. the mr13–mr11 thrust sheets were then passively translated on the trailing lower segment of mr10. moserende stage 3. initial imbrication of the mr08– mr05 thrust sheets resulted in an accumulated displacement of c. 200 m. the ramping was rooted in the 40 m décollement level along which the main translation of the trailing-end thrust sheets of the moserende section took place. the thrusting involved a complex relationship between mr07 and mr06 that may be interpreted as a connecting splay duplex (mitra & sussman 1997). above the l/r-unconformity, the deposits of the rubjerg knude formation probably reached a thickness of 20 m. moserende stage 4. the frontal part of the section was activated by c. 40 m translation of mr1 along the 30 m 149 décollement level over the lowermost trailing-end segments in the stortorn section. mr02 and mr03 followed this translation, whereas mr04 ramped up one level from the 40 m décollement level to the 30 m flat level that resulted in the initial ramping of mr04 up over the rubjerg knude formation on the back of mr03. the continued displacement consequently reorientated thrust sheets mr05–mr07 into more steeply dipping orientations. the trailing-end thrust sheets from mr08 and northwards were translated passively during this displacement. moserende stage 5. the frontal displacement of mr01 continued along the 20 m flat level over the trailingend segment of the stortorn section. mr02 ramped up along the footwall ramp at the trailing end of mr01 during a fault-bend rotation, which also included the lower segment of mr01u. a vertical thrust separation of c. 10 m brought mr02 up along the northern termination of the mr01 piggyback basin. during the passage of two intermediate ramps, an irregular anticline formed on the back of mr02 that had significant implications for the synsedimentary structures formed in the mr02 piggyback basin (see description of the moserende section, above). during mr04 thrusting, mr03 was imbricated along the upper 10 m flat level and the mr03b and mr03c thrust segments started to break through the piggyback basin. from the rear, mr04 was pushed by mr05 which had to pass up over the fault-bend-folded segment mr04u. together with mr06 and mr07, the mr05 thrust sheet moved up to the highest level indicated by the l/r-unconformity, situated c. 20 m above sea level on the back of mr06 and mr07, and their thrust faults were steepened into a nearly vertical position. moserende stage 6. in the frontal part of the moserende section, mr01 picked up a lower segment and thrust up to the 20 m flat level, which consequently also elevated the piggyback basin up into its present high level. the trailing-end ramp of mr01 formed the footwall ramp for the mr02 thrusting, which resulted in a fault bend of mr02 as well as mr03. this was followed by the final displacement of 18 m along the leading mr03 thrust. minor adjustments and re-orientation of mr04–mr07 followed the ramping of mr03, and the trailing-end thrust sheets mr08–mr13 were passively displaced by translation along the 40 m décollement level. moserende stage 7. during this stage, a complex duplex was formed by thrusting of the frontal lower segments, which also including the trailing-end segments of the stortorn section. moserende stage 8. the final displacement along the leading thrust-fault ramp in the moserende section progressed up along the stortorn trailing-end footwall ramp. moreover, the fault-bend folding due to thrusting in the stortorn section brought the thrust sheets into their present steeply dipping orientation. moserende section: summary data balanced length (l0): 1120 m cross-section length (l1): 650 m shortening (δl): 470 m compression: 40.2% stortorn section the most important development in the stortorn section was the change from the lowermost 40 m décollement level to the 30 m décollement level. the ramp, or progressive development of lower ramps, which marked the change, is here referred to as the stortorn lower segment footwall ramp, and was located somewhere near the thrust between st04 and st03. thus, the st03–st01 thrust sheets had their lower décollement level at 30 m, whereas the thrust faults related to st04–st10 were rooted in the 40 m décollement level. formation of a duplex complex comprising the lowermost thrust segments exposed the stortorn formation, the oldest strata involved in the thrusting. in the frontal part of the section, a complex stacking of lower segments, remaining in the subsurface from displacement in the grønne rende section, resulted in duplex formation that elevated st01–st03 about 30 m above the reference level. due to arguments presented later (see grønne rende section) the duplex stacking had to have been contemporaneous with the shortening of the grønne rende section. in the stortorn section, seven stages have been differentiated of which stages 1–5 are illustrated in fig. 113. stortorn stage 1. this stage is a direct continuation of the displacement in moserende stage 4. in the stortorn section, deformation was initiated by imbrication of st07, st09 and st10 with an accumulated displacement of about 50 m. this resulted in a ramping 150 skærumhede group active thrust fault rubjerg knude formation moserende 2 mr13 mr12 mr11 m mk1 moserende 3 mk1 mr13 mr12 mr11 moserende 4 mk1 mr moserende 5 mk1 mr13 moserende 6 mk1100 m n s moserende 1 mr13 mr12 mr11 mr10 mk1 fig. 112. dynamic model of progressive deformation in the moserende section illustrated in six sequential restoration cross-sections. the six stages demonstrate steps in the development between the balanced cross-section and the structural cross-section (plate 2); thus the starting and final positions are not shown. the red lines indicate the active displacement surfaces in each deformation stage. the basic décollement sur face is the 40 m flat level. from this, the flat levels rise by 10 m onto the reference level (l/runconformity) defined as the 0-level. note (1) that the final two stages (7, 8) discussed in the text are not illustrated, and (2) that the thrustsheet terminology in figs 112–122 is simplified (i.e. mr 3 on fig. 112 is equivalent to mr03 in the text). 151 stortorn footwall ramp stortorn footwall ramp stortorn footwall ramp mr10 mr9 mr8 mr7 mr6 mr5 mr4 mr1 st10 mr10 mr9 mr8 mr7 mr6 mr5 mr1 st10 mr1 st10 r13 mr12 mr11 mr10 mr9 mr8 mr7 mr6 mr5 mr4 mr3 mr2 stortorn footwall ramps 3 mr12 mr11 mr10 mr9 mr8 mr7 mr6 mr5 mr4 mr4u mr3a mr3u mr3c mr3b mr2 mr1a mr1u mr1 st10 st10 stortorn footwall ramps mr13 mr12 mr11 mr10 mr7 mr6 mr5 mr4 mr3c mr3b mr3a mr2 mr1 st10 st10u st10u mr8 mr9 stortorn footwall ramp mr9 mr8 st10 mr1 152 100 m n s skærumhede group active thrust fault rubjerg knude formation mr1 st10 st9 st8 st7 st6 stortorn 2 st5 mr1 mr1 st10u st10 st9 st8 st7 st6 st5ust8ust8ust9u stortorn 1 stortorn footwall ramp st10mr1 st9 st6 st7 st8 stortorn 4 stortorn footwall ramps stortorn 5 stortorn footwall ramps st10 mr1 st10 st9 st8 stortorn 3 stortorn footwall ramps fig. 113. dynamic model of progressive deformation in the stortorn section illustrated in five sequential restoration cross-sections. the cross-sections demonstrate steps in the development between the balanced crosssection and the structural cross-section (plate 2); thus the starting and final positions are not shown. the red lines indicate the active displacement sur faces in each deformation stage. note that the duplex segments gr u refer to elements that had to be deformed contemporaneously with the shortening taking place in the grønne rende section; duplex segments st s refer to horse/splint segments. the final two stages (6, 7) discussed in the text are not illustrated. 153 st4 rf6 st1s st1st2 st3 st4 st4s st4s st4s st4s st2s st1s stortorn lower ramp rubjerg knude fyr ramp frontal footwall ramp st5 st5 st4 st2st3 rubjerg knude fyr footwall ramp rubjerg knude fyr footwall ramp stortorn lower segment footwall ramps rubjerg knude fyr footwall ramp stortorn lower segment footwall ramps st9 st7 st6 st5 st4 st3 st2 st1 rf6gru gru gru st8 st6 st5 st4 st3 st2 st1 st2s rf6 stortorn lower footwall ramp rf6 st1s st1 st1s rf6 154 up of st10 from décollement level 40 m to flat level 30 m, along which the translation displacement took place. most of the st09 thrust sheet was also ramped up by the formation of a lower duplex structure. both st10 and st09 were affected by fault-bend folding, which created a major distortion of the l/r-unconformity surface in the uppermost part of the thrust sheets. stortorn stage 2. during an accumulated displacement of about 200 m related to the st10 and st09 thrusts, ramping progressed with development of the first imbrications of st08 and st06. due to ramping from the lowest décollement level to the 30 m flat level in the trailing end of st06, a fault bend affected the st07– st10 thrust sheets that were translated piggyback on the st06 thrust sheet. this contributed to the steepening up of the st07–st10 thrust structures. in the frontal part of the section, the imbricate thrusting was initiated at st01–st03. accumulation of sediments referred to the rubjerg knude formation reached a maximum thickness of about 20 m, notably in the synformal troughs of st03 and st09 that formed during the progress over the ramps below. stortorn stage 3. at this stage, st05 was thrust about 40 m up over the footwall ramp on the back of st04. the st05 thrust was rooted in the 30 m flat level, and during a passage of a lower ramp from flat level 30 m to 20 m, the initial fault-bend-fold resulted in undulation of the l/r-unconformity at the top of the st05 thrust sheet. the st06 thrust sheet progressed over the footwall flat of st05, and both thrust faults were rooted down to the 30 m flat level along which the main translation of the sheets emplaced piggyback on st06 took place. the st08 thrust sheet was finally displaced along the upper flat at the top of the st07 piggyback basin. consequently, most of the 20 m thick succession in this piggyback basin was preserved and indicates the maximum level of sediment accumulation in the rubjerg knude formation during stage 3. thrusting of the st08 sheet along the footwall ramp on the back of st07 resulted in a further steepening of st09 and st10, while the frontal elevated parts of the st08–st09 thrust sheets became subject to erosion. the trailing-end lower segments of st10–st07 were over-thrust by the frontal parts of mr01 and mr02, corresponding to stage 7 in the moserende section. the accumulated displacement in stortorn stage 3 was of the order of 320 m. stortorn stage 4. during this stage, the st04 thrust sheet was thrust 40 m over the piggyback basin of st03, and st05 was thrust about 70 m over the upper flat on top of the piggyback basin of st04. during this relatively large displacement of st05, two lower duplex segments were picked up from the lower 40 m décollement level. after ramping over the stortorn lower ramp, the duplex segments participated in the thrusting up along the footwall ramp on the back of st04. the lower trailing-end segments of the stortorn section were finally thrust up along the steep footwall ramp on the back of st10 and subsequently the frontal parts of the moserende section were brought into their present upright orientation. erosion and redeposition affected the piggyback basins on st05 and st08, whereas thrusting over st07 and st04 sealed these piggyback basins. the accumulated displacement reached about 410 m. stortorn stage 5. a substantial displacement, in the order of 80 m, took place along the leading thrust in the stortorn section at this relatively late stage of development of the structures at stortorn. however, this is only a small amount of the accumulated displacement (c. 500 m) which is of the same order of magnitude as that taken up by the duplex stacking of the lower trailing-end segments of the rubjerg knude fyr and grønne rende sections. the ramping and thrusting of st01–st04 over this duplex structure explains the high elevation of the l/r-unconformity and overlying piggyback basins in the frontal part of the stortorn section. the formation of the duplex stack comprising the lower duplex segments annotated gru in fig. 113 would have taken place only after the imbricate thrusting in the grønne rende section developed (see below). the combination of displacement at the leading edge in one section and stacking of lower duplex segments in another, indicates a continuous progressive thrust-fault evolution. during the propagation of st05, the trailing end of st04 was involved in a duplex formation that resulted in fault-bend folding of the earlier formed st05 lower duplex at the stortorn lower ramp. the piggyback basin on the back of the st05 thrust sheet was deformed into a north-verging syncline due to steepening. a similar re-orientation is seen in the thrustisolated piggyback basins in st10 and st09. a marked diapirism and remobilisation of mud in the st01–st03, st05–st07 and st09 thrust sheets suggests that the diapirism was related to the intensity of ramping, especially when the ramping involved the 155 fig. 114. dynamic model of progressive deformation in the rubjerg knude fyr section illustrated in four sequential restoration cross-sections. the cross-sections demonstrate five stages in the development between the balanced cross-section and the structural cross-section (plate 2). the red lines indicate the active displacement surfaces in each deformation stage. note that the tip of the rf04 thrust sheet was displaced by normal faulting during syntectonic deposition in the rf03/rf04 piggyback basin. lower level segments and fault-bend folding of these segments. stortorn stage 6. the final ramping of lower segments from décollement level 40 m to flat level 30 m at the base of st05 terminated the translation along the lowermost 40 m décollement level. for the sections further south, the lower décollement level was situated at the 30 m level. in the frontal part of the section, continued minor compression steepened the thrust structures, and the tips of st02 and st03 were eroded and deposited in the piggyback basin of st01. stortorn stage 7. the structural complex, including the 156 st01 thrust sheet and the underlying duplex structure, became fault-bend-folded during the thrust propagation related to the progressive deformation in the rubjerg knude fyr section. stortorn section: summary data balanced length (l0): 1125 m cross-section length (l1): 570 m shortening (δl): 555 m compression: 49.3% rubjerg knude fyr section the rubjerg knude fyr section roots into the 30 m décollement level. the most striking features developed in the rubjerg knude fyr section are the large olistoliths in the piggyback basin that were derived from the collapse and gravity gliding of a projecting segment of st04. five stages in dynamic development have been distinguished, which are illustrated by four cross-sections in fig. 114. rubjerg knude fyr stage 1. during sedimentation of the first 10 m of sand of the rubjerg knude formation, the rf05, rf04 and rf03 thrust sheets were thrust up along their footwall ramps. rf04 was displaced 90 m along the upper flat level (10 m level) before the frontal part propagated up along the upper ramp. with a displacement of about 30 m, this brought the nose of the rf04 thrust sheet up into the open air, above the sedimentation level of the rubjerg knude formation. the displacement on the other two thrusts amounted to c. 20 m, implying an accumulated displacement of 70 m. rubjerg knude fyr stage 2. the exposed nose of the rf04 thrust sheet slumped down along a normal fault into the piggyback basin of rf03. at the same time, the frontal nose of rf03 was eroded away and sedimentation of the rubjerg knude formation onlapped and covered these features. at the leading edge of the section, thrusting was initiated that brought rf01 and rf02 up over what was to become the trailingend segments of the grønne rende section. rubjerg knude fyr stage 3. the frontal imbrication of rf01 and rf02 progressed during sedimentation up to about 20 m above the main l/r-unconformity level. the rf05–rf06 thrust sheet ramped up onto the intermediate flat above the trailing-end segment of rf04. the rf04 thrust sheet was displaced about 70 m up along the relatively steep footwall ramp at the trailing end of rf03. due to the fault-bend folding of rf04, the rf05–rf06 hanging-wall ramp was rotated into a vertical position. rubjerg knude fyr stage 4. when the second ‘drop’ of the frontal part of thrust sheet rf04 took place, a c. 45 m long slab of the relatively thin thrust-sheet nose slumped down along a normal fault with a vertical separation of more than 10 m. the ‘drops’ may be regarded as two break-back sequences of the rf04 thrust sheet (in the terminology used by mitra & sussman 1997; see figs 99, 100). sediment accumulation continued in the piggyback basin to a thickness of more than 30 m, including the ‘dropped’ noses of rf04. the final accumulation in the piggyback basin took place while the displacement in the rubjerg knude fyr section was concluded more than 500 m laterally to the south. the translation progressed along the 20 m flat level on top of what was to become the lower trailing-end segments of the grønne rende section. rubjerg knude fyr stage 5. the continued displacement of rf04 resulted in structural propagation of this sheet above its own piggyback basin with the ‘dropped’ thrust noses. stage 5 in the rubjerg knude fyr section is interpreted to have been contemporaneous with stage 7 in the stortorn section in which compression brought the thrust sheets into their final, steeply inclined position. rubjerg knude fyr section: summary data balanced length (l 0 ): 525 m cross-section length (l 1 ): 260 m shortening (δl): 265 m compression: 50.5% grønne rende section the impressive imbricate fan composed of 12 upright thin thrust sheets is the essential element in the grønne rende section. as a consequence of the displacement in the imbricate fan, 550 m of trailing-end lower segments were left behind to be stacked in a duplex below the frontal part of the stortorn section (fig. 113). 157 four stages have been differentiated in the development of the grønne rende section, the first three of which are illustrated in fig. 115. grønne rende stage 1. the initial thrust-fault framework was a low-angle imbrication, about 20° on each upper hanging-wall ramp, which rooted down to the upper 10 m flat level. during thrusting, the upper thrust sheets were split up into three main segments with leading thrust faults below gr02, gr06 and gr11/ gr12 which ramped down to the main level of detachment in the 20 m flat level. the initial displacement of the imbricate fan is regarded to have been 20 m on each thrust. this implies that the accumulated displacement sums up to 240 m. gr01 was not affected by thrusting in the first stage, and 240 m of its lower trailing-end segment was consequently not displaced during this stage. the sediments of the rubjerg knude formation attained a maximum thickness of 15 to 20 m during this stage, with decreased thicknesses on the back of the gr06–gr08 thrust sheets, which were elevated to the highest position. grønne rende stage 2. the imbricate thrusting progressed with a displacement of 50 m on each thrust. this implies that the hanging-wall flats were fault-bendfolded while they passed the footwall ramps, resulting in a dramatic steepening of the thrust sheets. below gr10–gr12, the gr07u and gr08u lower segments formed a duplex structure that resulted in elevation and complex ramp-propagation folding of the sheets above. the accumulated displacement implies an increase in length of the trailing-end segment of gr01 in the order of 500 m, allowing for some adjustments due to the irregular duplex deformation. sediment thicknesses in the piggyback basin in the frontal part of the section increased to 25–30 m. grønne rende stage 3. finally, the leading-edge thrust was activated and gr01 was displaced 50 m up along its footwall ramp. the gr01 thrust roots in the lower 30 m décollement level, and the displacement of the hanging-wall flat up along the footwall ramp resulted in steepening of all the early-formed thrust elements (gr02–gr13). the displacements of the individual thrust sheets range between 60 and 70 m. the thrusting resulted in the final, almost vertical, orientation of the thrust sheets. in the rear part of the section, complex deformation of the duplex below gr10–gr13 was reflected in unusual folding of the beds in the gr13 thrust sheet where folds with horizontal axial planes were formed due to gravity collapse of the piggyback basins. grønne rende stage 4. this stage concluded the thrusting of the leading hanging-wall ramp-and-flat over the footwall ramp in the trailing end of the stenstue rende section and the subsequent final rotation of the gr02– gr05 thrust sheets. in the trailing end of the section, the rf01 and rf02 sheets concluded the displacement by thrusting from the trailing-end segments of gr12 up over the footwall ramp onto the back of gr13. moreover, gr13 was rotated into an upright position whereby the horizontal axial planes became vertically orientated (plate 1). grønne rende section: summary data balanced length (l0): 1080 m cross-section length (l1): 423 m shortening (δl): 657 m compression: 60.8% comment. the lengths are measured from the footwall ramp between rf01 and gr13 to the footwall ramp between gr01 and ss06, near the thrust truncation of the l/r-unconformity. stenstue rende section two markedly different structural complexes were formed during the development of the stenstue rende section. they were mainly caused by the displacement of the same thrust sheet (ss01) when it was displaced 200 m over the upper flat on top of the piggyback basin in the sandrende section. the frontal part of ss01 above the footwall flat of the sr04 thrust sheet is one of the complexes. the other structural complex is the chaotic breccia and gravity slumping in the northern part of the section that formed as the piggyback thrust sheets were transported over a minor antiformal stack in the central lower part of the section. the progressive dynamic development in the stenstue rende section is described in terms of five stages, the first four of which are illustrated in fig. 116. stenstue rende stage 1. four minor imbrications with an accumulated displacement of 70 m initiated the 158 development in the stenstue rende section. at the leading-edge thrust, a minor connection splay separated ss02 and ss03. most of the thrusting was located at the upper 10 m flat level for a distance of about 180 m, in the northern part of which it was eventually rooted down to the lower décollement level. the displacement of the ss05 thrust sheet followed the same system, but with a smaller translation along the upper 10 m flat level. at the trailing end, ss06 was thrust up along a steep footwall ramp, and here the formation of duplex structures was probably initiated. the thickness of sediment (rubjerg knude formation) that had accumulated by this stage amounted to 10 m. stenstue rende stage 2. the leading-edge thrusting shifted to the ss01 thrust sheet, which was displaced 20 m up along the footwall ramp (the trailing end of sr04 in the sandrende section). the ss01 thrust fault extended down via an intermediate ramp to the 20 m flat level, and about 200 m from the leading footwall ramp it stepped down the lower ramp to the 30 m décollement level. at the upper hinge of the lower ramp, ss01 was folded into a fault-bend anticline, a small detachment anticline. along the foreland-dipping limb of the anticline in the ss01 thrust sheet, a normal fault was formed that displaced the tip of the ss02 thrust sheet. furthermore, the ss03 thrust sheet fig. 115. dynamic model of progressive deformation in the grønne rende section illustrated in three sequential restoration cross-sections; the final stage (4) described in the text is not illustrated. the red lines indicate the active displacement surfaces in each deformation stage. note how the shortening due to the displacement along the 20 m flat level resulted in the substantial length of the ‘left over’ lower duplex segment between the 20 and 30 m flat level. 159 became steeply inclined, and the initial imbrication of the thin ss03 thrust sheet resulted in the separation of ss03 from ss04. in the trailing end of the stenstue rende section, duplex stacking of the lower segments in ss06 resulted in elevation of the l/r-unconformity more than 10 m above the mean level. the accumulated displacement ranged up to 160 m. stenstue rende stage 3. thrusting of ss04 progressed on the upper flat over the piggyback basin of ss03 with a frontal displacement of 80 m. the hangingwall flat of ss04 ramped up along the footwall ramp of ss03 and during this translation the nose of ss05 became fault-bend-folded into a syncline with a steeply dipping southern limb. the trailing end of the ss04 thrust sheet was translated along the 10 m flat level; it was pushed from the rear by the ramping of the trailing end of the ss05 thrust sheet whereby the ss06 thrust sheet also steepened up. sediment thicknesses in the piggyback basins increased to c. 20 m, and the accumulated displacement ranged up to 240 m. stenstue rende stage 4. the dramatic major foreland thrusting of the ss01 thrust sheet, which included about 200 m displacement of the hanging-wall ramp 160 fig. 116. dynamic model of progressive deformation in the stenstue rende section illustrated in four sequential restoration crosssections; the final stage (5) described in the text is not illustrated. the cross-sections demonstrate the development stages between the initial and final positions displayed in the balanced and the structural cross-sections in plate 2. the red lines indicate the active displacement sur faces in each deformation stage. 161 over the piggyback basin of the sandrende section, occurred contemporaneously with the formation of an antiformal stack above the trailing end of ss01. the creation of the antiformal stack had already been initiated by the earlier formation of the minor detachment anticline at the ramp splitting the lower segments of ss01 (the ss01u segments). a duplex duplication of the lower ss01u segments accentuated the anticline, and finally the ss03 thrust sheet riding piggyback on ss01 was folded into an anticline with a steep foreland-dipping southern limb (fig. 90). along this limb, a normal fault developed that displaced the frontal part of the ss04 thrust sheet. a chaotic soft sedimentary fault breccia was formed during the stretching and fault separation of ss04 (fig. 91). due to an extra push from the rear, the ss05 thrust sheet was displaced a further 30 m to the south, which resulted in the formation of a huge southerly overturned slump fold above the normal fault zone (fig. 89). the accumulated displacement totals about 470 m. the displacement of the ss01 hanging-wall flat up along the steeply dipping footwall ramp constrains the sequential thrusting of the stenstue rende relative to the sandrende thrusting. thus stage 4 could not have begun before the maximum sedimentation in the piggyback basin was accomplished in the sandrende section. the initial ss01 thrusting could be regarded as a growth fault, whereby the syntectonic accumulation of sand added to the steepening of the footwall ramp. the present vertical to northerly overturned orientation of the ss01 hanging-wall flat and ramp resulted from differential thrusting and fault-bend of the ss01u lower hanging-wall ramp. note also the re-orientation of the normal fault at the tip of ss02, which due to the same deformation was bent into a horizontal position. stenstue rende stage 5. this stage corresponds to stage 6 in the sandrende section, wherein the ss01 thrust sheet riding piggyback on sr04 was displaced by normal faulting (fig. 117, stage 6). stenstue rende section: summary data balanced length (l 0 ): 760 m cross-section length (l 1 ): 285 m shortening (δl): 485 m compression: 62.5% comment. the lengths are measured from the footwall ramp between gr01 and ss06 to the footwall ramp between ss01 and sr04. if the compression was calculated from the leading-edge thrust tip of ss01 to the trailing-end footwall ramp of ss06, l1 amounts to 455 m, δl = 305 m and the calculated compression would only be 40.1%. sandrende section the dynamic development of the sandrende section was formerly interpreted as a combination of diapirism and normal faulting caused by volume exchange during thrust propagation (sadolin et al.1997). the model presented here aims at an explanation of the development purely based on a thin-skinned thrustfault model including dif ferential ramping and duplex formation. thus, the diapirism is interpreted to be an effect of ramping and fault-bend folding growth, similar to the model of mitra & sussman (1997), but also including mud-mobilisation and exaggeration of backlimb thrusting. the normal faulting occurring in the sandrende section is interpreted as the effect of differential ramping of a lower trailing-end segment that created foreland-dipping features above a hangingwall ramp propagation along an intermediate footwall flat. six stages of dynamic development have been dif ferentiated in the sandrende section (fig. 117). sandrende stage 1. after initial deposition of a 3–5 m thick succession of rubjerg knude formation sediments, the sr04 thrust sheet started thrusting about 50 m over the upper flat. the dip of the footwall ramp was relatively gentle, only c. 14°, and in the 15 m flat level the thrust fault may be traced along a minor flat segment on top of the lower trailing-end segment of sr03 (sr03u). from the minor intermediate flat, the thrust fault rooted down to the 30 m décollement level along a 20° dipping footwall ramp of sr03u. note that an upper and lower sr04 hanging-wall ramp was introduced subsequently. sandrende stage 2. translation of the lower sr04 hanging-wall ramp along the intermediate flat established the anticline in the central part of the sr04 thrust sheet. the sr03 thrust sheet started to propagate towards its foreland along the upper 10 m hanging-wall flat, and the frontal part of sr03 was displaced 50 m over the upper flat on top of the piggyback basin of sr02. the tip of the sr02 thrust sheet propagated up along a growth-fault ramp, which caused the steeply dipping 162 163 orientation of the northern boundary of the piggyback basin at the top of the sr01 thrust sheet. the accumulated displacement ranged up to about 150 m, including the initial thrusting of sr01. sandrende stage 3. during stage 3, the thickness of the sediments of the rubjerg knude formation reached 20 m in the piggyback basins in the sandrende section. in the basin at the top of the sr04 thrust sheet, the thickness varied considerably. the reason for this variation is that the top of the anticline above the sr04 lower hanging-wall ramp was subjected to erosion while deposition continued in the frontal part, south of the anticline, as well as in the basin north of the anticline. on the foreland-dipping flank of the anticline, minor sets of normal growth faults governed sedimentation (fig. 87). the tip of the sr04 thrust sheet suffered minor erosion before deposition resumed during thrust propagation. this is documented by the angular onlap relationships described by sadolin et al. (1997). sandrende stage 4. the thrusting of sr04 continued with 50 m further displacement. below the trailing end of the sr04 thrust sheet, the sr03u lower segment was picked up and displaced onto the footwall ramp of sr02. this minor duplex and ramp thrusting accentuated the sr04 hanging-wall anticline, and normal faulting on the foreland-dipping limb progressed. above the crest of the sr03u detachment anticline, a significant normal fault complex developed. here in the sr04 thrust sheet, a dense network of conjugate normal faults (fig. 85) resulted from lateral extension due to flexural slip bend over the upper hinge of the lower footwall ramp. sandrende stage 5. the thrusting of sr01 propagated up along the lower and intermediate footwall ramp of the trailing segments of the brede rende section. during this ramping, the hanging-wall flat of sr02 progressed up over the piggyback basin of sr01. the tips of the sr03 and sr04 thrust sheets thus experienced fault bending up along the footwall flat of sr02. the atypical northerly overturned tip at the top of the sr02 sheet probably formed due to accentuated reverse faulting along a former established back-thrust. in the trailing part of the section, a minor satellite splay thrust developed, which broke through the sr04 thrust sheet from the hanging-wall flat to the footwall flat below ss01. sandrende stage 6. the final development of the sandrende section was dominated by complex duplex formation and fault-bend folding of the sr01 thrust sheet below the frontal part of sr02. during the thrust propagation over the footwall ramp of the trailing-end segments of the brede rende section, a fault-bendfolded syncline was formed in sr01, which resulted in normal fault displacement of the sr01 piggyback basin and the overlying frontal part of the sr02 thrust sheet. similar normal faulting affected the ss01 thrust sheet, which had over-thrust the piggyback basin on the back of sr04. due to the intense ramping and folding of sr01 and its underlying duplex (sr01u) into an antiformal stack, mud of the lønstrup klint formation was remobilised in sr01, which intruded through the hanging-wall flat of sr02 to form the diapir in the sandrende section. sandrende section: summary data balanced length (l 0 ): 775 m cross-section length (l 1 ): 440 m shortening (δl): 335 m compression: 43.2% comment. the lengths are measured from the footwall ramp between ss01 and sr04 to the footwall ramp between sr01 and br08, at the level where the ramps cut the l/r-unconformity. the volume lost in diapirism has not been considered, and a reduced amount of compression would result by measuring l 1 from the tip of the sr01 thrust sheet to the sr04 footwall ramp. brede rende section the development of normal faults associated with foreland-dipping features of hanging-wall ramps transfacing page: fig. 117. dynamic model of progressive deformation in the sandrende section illustrated by five restoration crosssections. note that stages 2 and 3 include syntectonic sedimentation of the rubjerg knude formation, mainly related to stage 2, and the thrust-fault configuration concluding stage 3. the cross-sections demonstrate the development stages between the initial and final positions displayed in the balanced and the structural cross-sections in plate 2. the red lines indicate the active displacement surfaces in each deformation stage. 164 lated along footwall flats has already been demonstrated in the previous sections. one of the best examples of such a normal fault relationship occurs in the brede rende section. an essential element for this development was the formation of a long thrust sheet, translated laterally more than 150 m along the upper flat. this is demonstrated by the seven stages of development recognised in the brede rende section, as illustrated by the five cross-sections in fig. 118. brede rende stage 1. the first stage differentiated here is the initial sedimentation of about 3–5 m of the rubjerg knude formation. this corresponds well with the thickness of sediments deposited initially above the l/r-unconformity in the sandrende section; this unit is considered to represent pre-thrust sedimentation, i.e. the sediment record prior to piggyback basin formation. brede rende stage 2. accepting that the thinnest preserved section of the rubjerg knude formation indicates the timing of the earliest thrusting, then thrusting in the brede rende section was initiated with the displacement of the br03 thrust sheet. the frontal part of br03 was displaced about 50 m over the upper flat corresponding to the relative foreland in front of the leading edge of thrusting. the br03 thrust fault probably rooted down to the 30 m décollement level. however, translation in the upper 10 m flat level cannot be excluded, and in this case the beds disturbed by hydrodynamic brecciation might be interpreted as thrust flats. in the trailing end of the section, the br06 hanging-wall ramp was the next thrust to break through and initiate the translation along the upper flat. brede rende stage 3. the br06 thrust sheet was further displaced c. 50 m over the upper flat. the trailing end of br06 was separated by a splay thrust at the footwall ramp, along which the br07 thrust sheet propagated contemporaneously with piggyback thrusting of br08. this stage is equivalent to the frontal thrusting during stage 5 in the sandrende section. brede rende stage 4. sediment accumulation in the piggyback basins increased up to about 15 m. the marked dif ference in thickness of deposits is clearly seen by comparing the br05 thrust sheet with the br06 thrust sheet. the roof of br05 was obviously capped at an earlier stage than br06 where sediments accumulated to more than twice the thickness of that in br05. brede rende stage 5. with a displacement of about 60 m, the br05 thrust sheet propagated up along the footwall ramp of br04 and onto the upper flat on top of the br04 thrust sheet. translation of the br06 thrust sheet progressed c. 60 m along the upper flat. the thrusting rooted down to the 20 m flat level on top of the trailing-end segment of the br03 thrust sheet (br03u). the accumulated displacement amounted to 150 m, including the ramping and translation of the br07 thrust sheet along the same 20 m flat level. brede rende stage 6. after the thrusting of br05 and br06 ceased, the br04 thrust sheet was translated c. 80 m. the br04 thrust fault included three ramps: an upper gently dipping ramp from the upper flat to the 5–10 m flat level, an intermediate ramp-bend of the br03 thrust sheet due to the presence of the formerly established br02 footwall ramp, and a lower ramp from the 10 m to the 20 m flat level. the translation of the br04 lower hanging-wall ramp along the footwall flat of br03 created the foreland-dipping bend that, combined with the bend due to the br03 ramping, formed a syncline in front of the br04 ramp anticline. the normal fault created parallel to the foreland-dipping features displaced the tip of the br06 thrust sheet. the vertical offset on the normal fault amounted to c. 20 m, which also included the displacement caused by the offset in front of the br05 thrust tip. brede rende stage 7. finally, the leading-edge thrusting of the section propagated over the trailing end of the kramrende section. above the footwall ramp of br01, a minor antiformal stack was formed and subsequently an irregular duplex formation affected the br01 thrust sheet during the last stage of deformation in the brede rende section. this phase developed into diapirism that intruded towards the thrust fault between br01 and br02. facing page: fig. 118. dynamic model of progressive deformation in the brede rende section illustrated in five sequential restoration cross-sections. the cross-sections demonstrate seven stages in the development between the balanced cross-section and the structural cross-section (plate 2). the red lines indicate the active displacement surfaces in each deformation stage. note that significant normal faulting occurred in the brede rende section during stages 5 and 6 while the hanging-wall anticline in the middle part of the br04 thrust sheet was formed. 165 166 brede rende section: summary data balanced length (l0): 815 m cross-section length (l1): 440 m shortening (δl): 375 m compression: 46.0% kramrende section in the central part of the kramrende section, a major diapir developed during the progressive thrusting. the kramrende diapir was the most distally located diapir in the thin-skinned thrust-fault system indicating that a certain amount of ramp propagation from a deeper décollement level (at least 30 m flat level) was needed for macroscopic-scale diapirism. south of the kramrende section, the décollement level gradually changed to a shallower position and the intensity of ramping decreased. seven stages of dynamic development have been differentiated in the kramrende section, as illustrated in the five cross-sections in fig. 119. kramrende stage 1. the thrusting in the kramrende section was initiated with leading-edge propagation along the kr01 thrust fault, which constituted an upper footwall ramp with a dip of 10°, a minor intermediate flat at the 15 m flat level, and a c. 15° dipping lower ramp connecting the thrust fault to the 30 m décollement level. the displacement was in the order of 100 m along the upper flat, where almost no sedimentation of the rubjerg knude formation took place. kramrende stage 2. subsequent to the early stage thrusting, the lowermost 10 m of the rubjerg knude formation was deposited; the sediment thickness in the kr01 piggyback basin was probably a little less. kramrende stage 3. the kr01 thrusting progressed about 60 m over the upper footwall flat of what was to become the mb04 thrust sheet, and the trailing end of the kr01 thrust sheet was elevated to the 15 m flat level by ramp propagation over the lower footwall ramp of mb04. a small duplex segment (kr01s) under the middle part of the kr01 thrust sheet was picked up in the thrusting and displaced to the upper footwall ramp hinge, where it formed a minor angular anticline. in the syncline between the anticline and the footwall ramp of kr01, the thickness of piggyback basin sediment accumulation increased to about 15 m before the kr02 thrust sheet propagated c. 50 m up along the ramp, and the kr02 hanging-wall ramp partly capped the kr01 piggyback basin. the accumulated displacement ranged up to about 260 m. kramrende stage 4. thrusting of the kr03 thrust sheet was initiated up along the northerly dipping footwall flat of kr02. the kr03 thrust fault included an upper and a lower relatively steep (c. 23°) ramp. the top of the kr02 thrust sheet was probably exposed to erosion, and the rubjerg knude formation is thus missing in this part of the section. the piggyback sediment pile increased to a thickness of 20 m, as indicated by the sedimentary section preserved above the l/r-unconformity at the top of the kr04 thrust sheet. from the trailing end of the kr01 thrust sheet, diapirism intruded through the footwall ramp and irregular mud diapirism developed in the kr02 thrust sheet. kramrende stage 5. with a displacement of c. 30 m, kr03 thrusting propagated over the two ramps that resulted in the fault-bend folding of two anticlines separated by an intervening syncline. kramrende stage 6. the kr04 thrust sheet was thrust over the fault-bend-folds formed in stage 5, simultaneously with limited continuation of kr03 thrusting. minor irregular duplex formation started to develop into mud mobilisation at the trailing end of the kr02 and kr04 thrust sheets. kramrende stage 7. the final thrust propagation of the kr03 thrust sheet concluded with a displacement of 30 m up along the footwall flat of kr02. at the bend between the footwall flat and the footwall ramp of kr02, a remarkable set of reverse faults developed (fig. 72). the kr04 thrust sheet, carried piggyback on kr03, was also displaced by the reverse faulting, a fact that testifies to the relative timing of kr04 piggyback thrusting and kr03 ramp propagation. the reverse faults are regarded as back-limb thrusts similar to the back-thrust features mentioned in stage 5 of the sandrende section. minor back-limb reverse faults facing page: fig. 119. dynamic model of progressive deformation in the kramrende section illustrated in five sequential restoration cross-sections. the cross-sections demonstrate seven development stages, of which stage 2 represents a purely depositional phase and stage 4 only includes minor displacement. the red lines indicate the active displacement surfaces in each deformation stage. 167 168 also developed at the crest of the fault-bend-folded kr01 thrust sheet. polyphase diapirism evolved in the trailing end of the kr03 and kr04 thrust sheets, such that the primary thrust-fault framework was partially destroyed. kramrende section: summary data balanced length (l0): c. 600 m cross-section length (l1): c. 300 m shortening (δl): c. 300 m compression: c. 50% comment. the lengths are measured from approximate positions on the footwall ramps bounding the kramrende section and the data must therefore be regarded as tentative estimates. martørv bakker section the development in the martørv bakker section was dominated by the translation of a thrust sheet that was more than 600 m long and only 20–30 m thick. during nearly 400 m of displacement towards the foreland, a lower segment transformed into a duplex that ramped at a relatively late stage and created a faultfig. 120. dynamic model of progressive deformation in the martørv bakker section illustrated in three sequential restoration cross-sections. the cross-sections demonstrate four stages in the development between the balanced cross-section and the structural cross-section (plate 2). the red lines indicate the active displacement sur faces in each deformation stage. note the significant depression formed in the hanging-wall block south of the martørv bakker normal fault. in this depression, diamictites interlayered with slump-slides were deposited. 169 bend-fold anticline and syncline pair. at the upper surface of the intervening limb between the fold pair, a foreland-dipping normal fault was formed, rather similar to the structural complex formed in the brede rende section. simultaneously with the sedimentation of a diamictite, three slump-slides filled the piggyback basin developed in a syncline created at the top of the hanging-wall block of the normal fault. the sequential restoration stages are illustrated in three cross-sections in fig. 120. martørv bakker stage 1. thrusting in the martørv bakker section started with foreland thrusting of mb02, and translation of the trailing-end duplex that constituted the kr01 thrust sheet emplaced piggyback on the mb04 thrust sheet, thrust up along the footwall ramp of mb03. the more than 600 m long mb02 thrust sheet was displaced c. 105 m over the footwall flat of mb01. the mb02 thrust fault included two ramps, an upper footwall ramp of mb01 and a lower ramp between the 20 m flat level and the 30 m décollement level. the lower ramp was located below the central part of the mb02 thrust sheet, where it acted as the final step for the décollement level change to the 20 m footwall flat level. it is thought unlikely that significant sedimentation occurred in the section during this stage. 170 martørv bakker stage 2. the mb01 thrust sheet was displaced about 100 m over the foreland of the stensnæs section along the leading-edge thrust. the mb01 thrust was rooted down to the 20 m flat level, and it can be traced further on to the 30 m décollement level by passing the central lower footwall ramp of mb02. minor adjustments along the hanging-wall flat resulted in formation of small duplexes along the thrust fault. in the trailing end of the section, the mb03 thrust sheet was thrust up over the footwall ramp of mb02, whereby an antiformal stack was formed due to the folding that also involved the mb04 and kr01 thrust sheets. at the base of the mb03 thrust sheet, the lower segments formed an irregular duplex, which accentuated the antiformal stack. the accumulated displacement ranged up to 290 m. martørv bakker stage 3. the final thrusting of the martørv bakker section was concluded by nearly 100 m displacement of the mb02 thrust sheet. the frontal hanging-wall ramp-and-flat was thrust over the piggyback basin of the sn04 thrust sheet in the stensnæs section. during the thrusting, the lower segment mb02u was activated and formed a fault-bend-folded lower duplex. above the hanging-wall ramp of the trailing-end segment (mb02u3), an anticline was formed at the surface of mb02 and a subsequent syncline above the hanging-wall/footwall flat became a piggyback basin. martørv bakker stage 4. the foreland-dipping limb of the fold pair at the top of the mb02 thrust sheet developed into a southerly dipping normal fault. the c. 10 m deep piggyback basin was filled with diamictitic deposits and slump-sheets that glided down from the top of the antiformal stack. deformation in the martørv bakker section concluded with the steepening up of the leading-edge thrust structures due to ramp bending in the stensnæs section. martørv bakker section: summary data balanced length (l 0 ): 1065 m cross-section length (l 1 ): 675 m shortening (δl): 390 m compression: 36.6% comment. the lengths are measured from the tip of the leading-edge hanging-wall ramp to the upper bend of the footwall ramp of the mb04 thrust sheet. stensnæs section in the stensnæs section, a number of conspicuous flexural slip folds occur which are interpreted to have resulted from the deformation that accompanied sequential footwall ramp collapse and subsequent ramp displacement of minor duplexes. it is significant that they occur in relation to the final ramping from the lower 20 m flat level to the upper 10 m flat level. eight stages have been differentiated in the development of the stensnæs section, of which stages 2, 4 and 6–8 are illustrated by the cross-sections in fig. 121. stensnæs stage 1. in contrast to the martørv bakker section, an initial sediment thickness of 5 m of the rubjerg knude formation is thought to have covered the stensnæs section. it should be noted, however, that typical lønstrup formation facies grade upwards into typical rubjerg knude formation facies in this distal part of the rubjerg knude glaciotectonic complex; the l/r-unconformity is not clearly developed, and location of the formation boundary can be difficult. the affinities of the sediment packet referred to above are thus debatable. stensnæs stage 2. thrusting in the stensnæs section was initiated with c. 100 m displacement of the sn02 thrust sheet over the upper flat. the thrust fault ramped down to the 10 m flat level, which separated the upper and lower segments of the sn04 thrust sheet, simultaneously with stacking the sn03 thrust sheet into a northerly dipping duplex complex along the footwall ramp of sn01. stensnæs stage 3. accumulation of the rubjerg knude formation increased to a sediment thickness of 10 m. sedimentation was restricted to the piggyback basin of the sn04 thrust sheet, as well as on the foreland south of the frontal tip of the sn02 thrust sheet. stensnæs stage 4. the piggyback basin on the back of sn04 was sealed in by the overthrusting of the mb02 thrust sheet; this is equivalent to stage 3 in the martørv bakker section. stensnæs stage 5. as a trailing-end structural complex to the ulstrup section, the thrust sheets of the stensnæs section were translated together with the ul02 thrust sheet over the ulstrup footwall ramp onto the hanging-wall flat of the foreland. during ramping, the sn01 and sn03 thrust sheets were separated into small duplex segments. flexural-slip folding and polyphase 171 ul1 intermediate footwall rampsulstrup footwall ramp stensnæs 2 ulstrup footwall ramp mb1 sn4 sn4u sn4u sn1 ul2sn1 sn3 sn2 stensnæs 4 ulstrup footwall ramp mb2 mb1 sn4u sn4u sn4 sn3 sn2 sn1 sn1u ul2 ? stensnæs 6 mb1 sn4u sn4 sn3 sn2 sn 4u sn1 sn1u ul2 ul1 mb2 ? ? stensnæs 7 & 8 stensnæs ramp mb2 mb4 sn4 sn3 sn2 sn1 sn1u ul2 ul1 lønstrup klint formation active thrust fault rubjerg knude formation 100 m n s fig. 121. dynamic model of progressive deformation in the stensnæs section illustrated in four sequential restoration cross-sections. the cross-sections demonstrate five of the eight stages in the development described in the text between the balanced cross-section and the structural cross-section (plate 2). the red lines indicate the active displacement surfaces in each deformation stage. hydrodynamic brecciation resulted from the ramping (figs 53, 57, 58). the accumulated displacement ranged up to 35 m, whereas the length of the hanging-wall flat in the 10 m flat level amounted to 500 m. at the leading edge of thrusting, the ul02 thrust sheet initiated the thrusting up over a stepwise ramp. stensnæs stage 6. during this stage, about 10 m of the rubjerg knude formation was deposited in the piggyback basin at the top of the sn02 thrust sheet. the sedimentation level was probably up to 20 m above the l/r-unconformity, inferred from the elevated position of the sn02 thrust sheet. however, this is uncertain and the sediments were either never deposited or eroded away during later thrust elevation. in the northern part of the section, the sn04 thrust sheet propagated up along the footwall ramp of the earlier created sn02–sn03 duplex. this resulted in fault-bend folding of the sn04 thrust sheet and its piggyback basin as well as the overlying mb02 thrust sheet. this 172 stage correlates with stages 3–6 in the martørv bakker section. stensnæs stage 7. the sn01 thrust sheet was displaced about 50 m over its lower segment (sn01u), and together they were thrust onto the footwall ramp-and-flat of the ul02 thrust sheet. during the thrust-fault propagation of the ul02 thrust sheet over the footwall ramp of ul01, the sn01 and sn02 thrust sheets, piggyback translated on ul02, were bent into c. 30° dipping position. finally, the sn04 thrust sheet was displaced up along the footwall ramp of sn03 during dif ferential duplex formation along the sn04 hanging-wall ramp. stensnæs stage 8. the frontal parts of the sn01, sn02 and sn03 thrust sheets, as well as the anticlinal crest of the ul02 thrust sheet (formed above the upper hinge of footwall ramp of ul01), were significantly eroded, and a local piggyback basin was formed above the transition between the stensnæs and ulstrup sections. to the north of this piggyback basin, the elevated and exposed tips of the sn02–sn04 thrust sheets gravity-slumped out into the basin, where they were deposited as olistoliths, 1–5 m in size. stensnæs section: summary data balanced length (l0): 350 m cross-section length (l1): 180 m shortening (δl): 170 m compression: 48.6% 173 ulstrup section thin-skinned thrusting in the ulstrup section involved the remarkable translation of extensive, thin thrust sheets over the footwall flat of the foreland. cohesion of the thrust sheet was probably increased by ground frost in the upper part of the thrust sheet, while the hanging-wall ramp-and-flat slid on a thin zone of mobilised mud. during translation, piggyback sedimentation varied considerably. six stages have been differentiated in the development of the ulstrup section; stages 1–3 and 5 are illustrated by the cross-sections in fig. 122 (see also fig. 121). ulstrup stage 1. thrusting in the ulstrup section initiated with frontal ramping of the ul02 thrust sheet over a two-stepped footwall ramp of what was to become the ul01 thrust sheet. this ramping resulted in the formation of two, fault-propagating folded anticlines, which were separated by a shallow, broad syncline. the leading edge of the ul02 hanging-wall ramp was displaced about 25 m over the c. 5–10 m thick rubjerg knude formation deposited in the foreland (at the top of ul01). the ul02 hanging-wall flat extended along the upper 10 m flat level for about 400 m, terminating to the north at the foreland footwall ramp rooting down to the 20 m décollement level. thrust propagation up over this ramp formed a hanging-wall anticline at the trailing end of the ul02 thrust sheet. between the anticline at the trailing end and the anticline at the upper footwall ramp of ul01, a piggyback basin formed in which glaciolacustrine sediments were deposited to form the small, ephemeral ulstrup lake. fig. 122. dynamic model of progressive deformation in the ulstrup section illustrated in four sequential restoration crosssections. the cross-sections demonstrate four of the eight stages in the development described in the text between the balanced cross-section and the structural cross-section (plate 2). the red lines indicate the active displacement surfaces in each deformation stage. 174 ulstrup stage 2. thrusting along the ul02 hanging-wall ramp-and-flat progressed with an accumulated displacement of 230 m. the glaciolacustrine deposits of the ephemeral ulstrup lake participated in the ramp-propagating-folding. during translation along the upper 10 m flat level, the trailing end of the ul02 thrust sheet was probably covered by sediments, which subsequently became eroded. this event in the rubjerg knude formation corresponded to stage 6 in the stensnæs section. ulstrup stage 3. the long lateral thrusting of the ul02 thrust sheet along the upper flat resulted in 550 m of displacement, and at the lower trailing end, the hanging-wall ramp became detached to the upper footwall ramp of ul01. above this ramp, conspicuous flexuralslip folds, similar to the folds developed in the stensnæs section, were formed in the ul02 thrust sheet. ulstrup stage 4. glaciofluvial sands were deposited upon an erosional surface capping the ulstrup lake sediments (all rubjerg knude formation). depositional base level was probably equivalent to that experienced in stage 8 in the adjacent stensnæs section (see above). ulstrup stage 5. the final foreland thrusting took place as the ul01 thrust sheet was thrust over the upper ramp of the foreland and propagated about 200 m to the south. when the anticline above the ul02 hanging-wall ramp approached the footwall ramp of the foreland, where a fault-bend formed continuously during the propagation of the ul01 thrust sheet, a narrow channel was formed in which coarse-grained glaciofluvial gravel was deposited (figs 27, 122). the gravel also included redeposited frozen blocks of sand, testifying to the ground-frozen conditions of the environment (fig. 28). ulstrup stage 6. at the leading edge of the ul01 thrust fault, the deformation concluded with 150 m of displacement over the upper footwall flat of the foreland. during the translation of the ul01 thrust sheet over a minor depression in the foreland, a sandy mud volcano developed due to trapping of the high water pressure close to the leading-edge thrust. the sediment extrusion resulted in chaotic disturbances in the central part of ul01. ulstrup stage 7. the last stage of development in the ulstrup section involved sedimentation of the uppermost post-tectonic deposits of the rubjerg knude formation. the conglomerate (of stage 5) was covered by sand, and deposition in the foreland covered the leading-edge thrust at tvonnet rende (for location see plate 1). ulstrup section: summary data balanced length (l0): c. 1350 m cross-section length (l1): c. 850 m shortening (δl): c. 500 m compression: c. 37% summary of dynamic development the dynamic development of the complex is summarised in figs 123 and 124. from the scheme in fig. 123, it is clear that the rubjerg knude glaciotectonic complex developed in sequential progressive stages during syntectonic sedimentation of the rubjerg knude formation. the thin-skinned thrust-fault complex developed mainly as piggyback thrusting with proximal thrust sheets being displaced contemporaneously with activation of the distal thrust fault. during the advance of the thrust-fault complex, the position of the décollement zone shifted progressively to deeper levels. the dynamic development can be summarised in eight steps that resulted in the formation of eight characteristic thrust-fault structure types (fig. 124). fig. 123. summary scheme of the syntectonic sedimentary development in the rubjerg knude glaciotectonic complex. the deformation stages for each section, as described in the text, are indicated here by the symbol #. 175 1. long lateral translation of a thin thrust sheet took place over the foreland. the ramp was rooted in the uppermost shallow décollement level 1 at a depth of c. 10 m from the top surface. 2. ramps became rooted in décollement level 2, and the increase in ramp height, amounting to about 20 m, is regarded to be the cause of the duplex folding at ramp collapse. 3. the hanging-wall anticlines became dominant structures with hinterland-dipping piggyback thrust sheets on the back limb. as the ramps extended down into décollement level 3, the height of the ramps increased and consequently the hangingwall anticlines increased in size. 4. the antiformal stack developed, which included long-distance translated piggyback thrust sheets that were folded in a hanging-wall anticline. in relation to the antiformal stack, foreland-dipping thrust faults occur that were accompanied by normal faults. 5. the prominent imbricate fan formed above décollement level 2. the initially gently to moderately dipping imbricated thrust sheets were re-orientated into steeply dipping positions due to lateral translation of the imbricate fan along décollement level 3. 6. this step involved the subsequent deformation of the lower duplex segment not incorporated in the imbricate fan. this lower duplex segment was imbricated and the sub-segments were displaced into a duplex stack during push from behind by a progressing hanging-wall ramp, rooting in décollement level 3. 7. this step involved differential duplex stacking and imbrication of thrust-fault sheets. the piggyback basins vary in elevation due to differences in duplex stacking. furthermore, the variation in duplex stacking reflects the shift from décollement level 3 to 4 (corresponding to a shift in the décollement surface from 30 to 40 m). 8. the fault-bend-folded duplex units were formed. the formation of these duplex units was only possible because the four thrust-fault flat levels had developed, and thus the duplexes could be stacked and subsequently fault-bend-folded during maximum compression and translation along décollement level 4 (fig. 124). the thickness of sediments that accumulated contemporaneously in tectonically correlated piggyback basins decreases from north to south. thus the depocentre was situated in front of the last activated thrust section, and the depocentre gradually shifted to a more and more distal position. correlation of the syntectonic progressive development of the complex shows that sedimentation was contemporaneous with thrusting rather than there being an alternation between periods of active thrust faulting and periods of deposition. furthermore, it indicates that the ice margin was not melting back during the formation of the complex but advanced in a continuous progressive gravity-spreading process. 176 ice fault-bend-folded duplex imbricated duplex stack duplex stack imbricate fan antiformal stack hanging-wall anticline duplex folds at ramp collapse foreland translated thrust sheet distance t im e sn fig. 124. model of thrust-fault structure types formed during the progressive deformation of the thin-skinned glaciotectonic thrust-fault complex. the model outlines a progressive development in eight steps resulting in the formation of eight characteristic thrust-fault structure types, the first to develop earliest and continuously in the distal part of the complex, and the last to be formed in the most proximal part of the complex (see text for details). 177 lønstrup klint formation rubjerg knude formation stortorn formation 178 discussion the observations that form the basis for the description of the structural geology, mechanical behaviour and dynamic development of the rubjerg knude glaciotectonic complex, raise important questions with respect to understanding the framework and nature of thin-skinned thrusting related to glacial deformation; seven topics have been selected for further discussion below. the basis for understanding a structural complex is to describe the tectonic architecture and the range of structures it contains from microscopic to macroscopic scale. the discussion of thrust-fault architecture leads to evaluation of the reliability of the balanced cross-section. consideration of thrust brecciation and diapirism leads naturally to a focus on the thrustfault dynamics, and the significance of the rate of deformation. the dynamics associated with the syntectonic deposits and the formation of piggyback basins merit discussion, as does the interpretation of a proglacial contra subglacial deformational setting. the final topic deals with the geological setting of the complex, including the timing of the event that created it. thrust-fault architecture a prerequisite for understanding the thrust-fault architecture is a familiarity with the terminology (see appendix 2). the macroscopic structures encountered in thin-skinned orogenic belts are all recognisable in the glaciotectonic complex. mesoscopic structures such as folds and faults are similarly recognisable. however, small-scale structures such as joints, cleavage and fabric are more difficult to recognise (except for hydrodynamic brecciation), and this may be one of the major dif ferences between soft sedimentary deformation and hard-rock deformation. it seems likely that joints and fractures in soft sediments would be able to re-heal after deformation. thus, a large number of minor reverse faults must have formed in thrust sheets during ramp propagation (fig. 67), but appear to have disappeared again after subsequent thrust sheet propagation along the flat, as they have not been observed with the exception of the in situ positions related to ramp bend (fig. 85). there is an approximation to a right-angle relationship between the footwall ramp and the back-thrust faults, which indicates that an increase in the dip of the footwall ramp results in a decrease in the dip, in the opposite direction, of the back thrust. moreover, a steeper and higher footwall ramp also corresponds to an increase in displacement along the back-thrust fault. thus one can regard the kr01 (kramrende) back-thrust faults as structures related to initial faulting in the progressive deformation (fig. 67), and the kr04 back-thrust splay faults as a structural element related to a developed phase of progressive thrust-fault deformation (fig. 72). major back thrusting at the back of sr02 (fig. 84) represents a mature phase in the progressive thrust faulting. the apparent lack of joints and fractures reflecting ramp propagation could probably be explained as having been absorbed in the hydrodynamic brecciation process. among the structural elements analysed during the interpretation of the balanced cross-section, the duplex structures create the most interesting problems. firstly, the interpretation of the duplex imbricates in the stensnæs section provides an explanation for the complicated fold framework. secondly, the interpretation of the duplex below the frontal part of the stortorn section links the hidden duplex segments at the base of the grønne rende section with the duplex stacking below st01–st03. thirdly, the normal faults can be interpreted to have been related to the ramping of lower duplex segments. if the normal faults are regarded as foreland-dipping faults or part of a foreland-dipping duplex, the model for duplex formation suggested by contreras & sutter (1997) may be relevant for the understanding of the foreland-dipping faults. in their model for formation of forelandor hinterland-dipping duplexes, they considered two factors: u = distance of displacement along the upper flat, and s = length of duplex segment. in a regime where the ratio u/s is greater than one (u/s > 1), foreland-dipping duplexes are formed; in a regime where u/s < ½, hinterland-dipping duplexes are formed. in regimes where ½ < u/s < 1 or u/s = 1, antiformal stacks or angular antiformal stacks, respectively, are formed. this corresponds well to the interpretation presented here of the rubjerg knude cross-section, where most thrust sheets are displaced by less than their length, and consequently the main orientation of thrust sheets is hinterland dipping. according to the model of contreras & sutter (1997), foreland-dip179 ping duplexes are formed when a duplex segment is displaced along an intermediate or upper flat for a distance equal to, or more than, its length. a consequence of this is that a roofing thrust sheet will be displaced in front of the foreland-dipping upper footwall flat, where normal faulting will take place. this is interpreted to be the case for the normal faults in the martørv bakker and brede rende sections (see fig. 66). the normal fault developed in the stenstue rende section may also be regarded as an expression of the latter regime in the suggested model. a foreland-dipping feature may well reflect the foreland-dipping limb of a hanging-wall anticline formed above a laterally displaced hanging-wall ramp. however, the normal fault-displaced thrust sheet must be thrust over the duplex segment before it was thrust faulted together with its roofing thrust sheet, and then as translation continued attached to the hanging-wall flat until the displacement was concluded by the normal faulting over the tip of the duplex segment. the consideration of duplex formation naturally leads to a focus on the changes in décollement levels. when a duplex segment is formed, there would normally be an early décollement surface at a shallow level, succeeded by a shift to a deeper level connected with a new footwall ramp. the upper décollement level (the 10 m level or corresponding gently dipping ramp to the leading edge) was probably the first to be formed in the proximal part of the complex and probably also the last to form in the distal part (fig. 124). from the cross-section, it is indicated that the 10 m décollement level extended for about 1 km, but ended up with a distance of only 600 m. the 20 m décollement level was the next to take over, and during the establishment of a related ramp, footwall ramp imbrication progressed, modifying the ramp transition connecting the two décollement levels. the length of this décollement level might have been of the same scale, but only c. 400 m is preserved as a lower footwall flat. the 30 m flat level is the dominant décollement level extending from the middle of the martørv bakker section to the middle part of the stortorn section, where finally the 40 m décollement level was developed. note that the present-day position of the lowermost décollement surface is at about 45 m b.s.l. due to the regional, very gentle dip to the north of the l/r-unconformity that serves as the reference level, corresponding to the 0 m flat level. from the cross-section, it can be seen that the dips of the ramps increase from gentle (5–15°) in the zone between the upper surface flat to the 10 m flat level, to 25° dips between the 10 and 20 m levels, and reaching up to 35° between the 20 and 30 m flat levels. the ramp dips with steeper angles in the cross-section, arise from over-steepening or superimposed tilting during ramp propagation (fig. 9). thus, in the proximal part of the complex, dips between 35° and 45° are interpreted as the primary dips of ramps rooting down to the deepest décollement level at 40 m, which is incorporated in the model for the fold-imbricate duplex units. the increasing dips of ramps are interpreted to have resulted from the increase in fracture angle as a function of increase in normal stress (change of levels) and increase in shear stress (increasing force required to move thrust sheets). this is implied from the shape of the mohrenvelope in the mohr diagram (hobbs et al. 1976), and it is suggested to be a basic relationship for glaciotectonic fracture and fault deformation (pedersen 1996). balanced cross-section in the balanced cross-section, the changes in dip angles are responsible for the insertion of a number of small triangular-shaped duplex segments, which are incorporated in the geometric construction and annotated as splints (horses) (plate 2). it is not known how many splints exist in reality. a few have been recognised as structural identities (kr01 s in kramrende), but it is likely that space deficits or excesses have been absorbed in mud-mobilisation or differential small-scale anastomosing fracturing. dif ferential fracturing and thrust-fault formation with a spacing of only 1 m has been documented in the moserende section (fig. 27), indicating that the duplex segmentation does exist. hence it is probable that a much more differential translation took place than is indicated in the crosssections of the dynamic model of thrust-fault propagation (fig. 112). the reliability of the approximations in construction inherent in a balanced cross-section is founded in the area balance. the main calculation of the balance indicates that the shortening amounts to approximately 50%, with l 0 = 12 km and l 1 = 6 km. the area of the l 0 cross-section (l 0 multiplied by stratigraphic thickness) amounts to 340 000 m2, and the area of the l 1 cross-section (l 1 multiplied by measured thickness of the retrodeformed cross-section) is 382 500 m2. the dif ference amounts to 11%, which is interpreted as a consequence of the erosion of the thrust sheets in the proximal part of the complex (grønne rende section – ribjerg section). the detailed calculations of areas 180 for the area balance are summarised in table 1, and documented in plates 2a and 2b. the amount of erosion indicated from the area balance differs markedly from the 80% erosion estimate by gry (1941), and supports the argument that gry’s cylindrical thrust-fault model was incorrect. considering the amount of erosion, the question arises: why is the preservation potential so great? three factors are suggested here to answer this: (1) the steeply orientated thrust sheets were partly packed by the sand fill in the piggyback basins, (2) the thrust-fault deformation resulted in a strain hardening that consolidated the complex, and (3) as the sole of the approaching ice sheet advanced across the proximal part of the complex, the over-pressured pore water migrated from the hanging-wall ramps and flats of the thrust sheets to the hanging-wall flat of the ice sheet, facilitating the over-thrusting of the footwall block which subsequently comprised the thrust-fault complex. thrust brecciation and diapirism in order for a thrust sheet to move, a fracture must be created that can develop into a plane of thrusting. the initial fracture is formed when the failure limit is reached in a system subjected to pressure (loading and lateral compression). a recurring question, and an apparent conflict in reasoning, is why fault planes develop, leaving the rest of the thrust sheet preserved? since the sedimentary unit forming a thrust sheet is subjected to the same amount of confining pressure, it might be expected that a muddy mass of collapsed sedimentary units just as well could have been the result? it is well known that an increase in pore-water pressure results in failure and initialisation of fractures along surfaces of anisotropy, as described for orogenic systems by hubbert & rubey (1959). however, in soft sedimentary deformation with lower confining pressure and smaller shear strength, as well as smaller coefficient of internal friction, the limits of fracture formation and complete collapse are much narrower. the structures developed in the ulstrup section reflect these conditions. the anastomosing jointing and mud mobilisation at the tip of the ul01 thrust sheet reflect the stage of near collapse (fig. 49). the thick zone of hydrodynamic brecciation along the hangingwall flat reflects the same tendency towards collapse, and speculations about the influence of ground-frozen conditions on the preservation of the thin thrust sheets during translation over the foreland are relevant. ground-frozen conditions are interpreted to have affected that part of the thrust sheets elevated above the ground surface; the freezing of the sediments in the thrust sheet may result in more brittle behaviour, whereby cracks formed (figs 44, 46). however, due to the high pore pressure maintained along the hanging-wall flat, the cracks were filled with sand pumped into the cracks by the over-pressured pore water from the base of the thrust sheet. hydrodynamic brecciation is evidently related to the hanging-wall ramp-and-flat. brecciation was initiated at an episedimentary stage with the formation of ball-and-pillow structures due to sediment loading. when the loading increased by over-thrusting, the ball-and-pillow formation progressed further and hydrodynamic brecciation was concentrated at the hanging-wall flat. small-scale mud diapirism took place, with chaotic folding developing into polydiapirs (figs 54, 55, 77, 78, 86, 88). polydiapirism and mud-mobilisation are considered to have developed simultaneously and with an increasing degree of disordering and size of diapir in progressive stages of deformation. many of the mesoscopic diapir structures recognised in the rubjerg knude glaciotectonic complex can be compared with the multi-wavelength gravity structures described in the model analysis by weinberg & schmeling (1992). the formation of large-scale diapirs is suggested to have been related to thrustfault deformation of a deep-seated hanging-wall flat that propagated up to surface level along a set of relatively steep footwall ramps. during propagation up along a lower ramp to an intermediate flat, and ramping from the 20 m intermediate flat level to the 10 m flat level, polysequential hanging-wall anticlines formed, and were subsequently destroyed by mudmobilisation initiated from the deep-seated thrust zone of the hanging-wall flat. some of the soft sedimentary xenoliths floating in the mud diapirs can be viewed as relicts of anticline crests (fig. 74). staircase-like ramp propagation is indicated for the kramrende diapir and the sandrende diapir, but is not so obvious in the case of the brede rende diapir. intrusive remobilised mud is evidently related to the footwall ramp propagated hanging-wall flat of gr01, and the mud mobilisation in the thrust sheets of the stortorn and moserende sections are all easily identified with sequential ramping from the deepest décollement level. 181 thrust-fault dynamics the difference in thrust-fault development that relates to the upper flat level (10 m) is very marked when the ulstrup section is compared to the grønne rende section. thus the foreland regime in the latest stage of deformation is characterised by thin, very long sheets subjected to horizontal translation over the footwall flat of the foreland. in contrast, the grønne rende section probably formed an imbricate complex of smaller, moderately dipping thrust sheets when this section was adjacent to the foreland. there is no obvious reason for this difference, although minor differences in lithology and dif ferences in environmental conditions (frozen or unfrozen ground) could be viewed as contributing factors. however, there is an invisible condition which must be considered, namely the velocity of deformation. at the initiation of any deformation, the velocity is zero; the velocity then increases until the displacement is brought to a halt at the edge of the foreland during decreasing velocity. fast deformation results in more fractures than slow deformation. it is therefore suggested that the imbricate structures in the central part of the thrust-fault complex were initiated during the fastest advance towards the foreland and that the long-distance translation of unbroken thrust sheets relates to decreasing velocity or slow advance. in a discussion of the velocity of thrust-fault propagation, the question of rates and timing is inevitable. the youngest dating of the stortorn formation is 30 000 years b.p., while the oldest dating of the rubjerg knude formation is about 29 000 years b.p. and the oldest dating of the ribjerg formation is 26 000 years b.p. thus, a time span of 3000 years is estimated for the calculated shortening of 6 km, which indicates an average velocity of 2 m per year. the peak velocity of the deformation must evidently have been more than 2 m per year, taking into account the acceleration and deceleration. however, the velocity would also have been much higher if deformation had progressed in periodic steps rather than continuously. a step-like process would have involved periods of no movement alternating with higher velocity in the periods of advance. with respect to the rubjerg knude glaciotectonic complex, the summary of the dynamic development suggests that a continuous progressive deformation process characterised the formation of the complex (fig. 123). although the developments of the sections are described separately above, the deformational overlap from one section to the next links the sections in a continuous dynamic development. syntectonic deposition the concept of piggyback basins was originally related to large-scale regional orogenic settings (ori & friend 1984; ricci lucchi 1986). however, as applied here the term is used for the syntectonic deposits of the rubjerg knude formation that were laid down in basins structurally overlying moving thrust sheets. the initial depositional environment of the rubjerg knude formation was a relatively flat lowland, dominated by shallow lakes in an outwash plain bounded by an ice margin to the north. the plain was probably gently dipping towards the north due to isostatic loading of the ice cap. judging from the variation in thickness of the rubjerg knude formation (30 m in the proximal part to only about 10 m in the distal), the dip of the plain was not more than 2°. as the thrust belt propagated southwards, the plain became separated into smaller, more or less isolated basins characterised by steep slopes and uneven relief. the most distinctive deposits in these basins are the sedimentary breccias and slumped thrust sheets derived from the tips of up-thrust thrust sheets. three types of syntectonic slump/slide deposits can be differentiated. the first type involves deposition of coarse clasts up to metre size, which were rotated indicating transport as sedimentary clasts enveloped by sandy mud. this deposit type is regarded as being related to the distal part of the thrust-fault system, and is exemplified by the piggyback basin in the stensnæs section (fig. 56). the second type is characterised by isoclinally folded slump sheets interlayered with matrix-supported coarse clastic diamictite. this indicates that the source was very close to the depocentre, although the slump sheets were detached from their roots and were transported independently by gravity gliding into the basin. the piggyback basin in the martørv bakker section represents this deposit type (figs 23, 64). the third type comprises slump-folded sheet segments that can be traced directly, or correlated over short distances, back to the source of the thrust sheet; this type is regarded as being related to the proximal part of the system. the next step in the development would be that of thrust sheets displaced by normal faulting, but lacking depositional features such as sedimentary breccias. however, this type of dynamic development is strictly tectonic. the major 182 slump fold occurring in the stenstue rende section (fig. 89) may be regarded as a transition from a sedimentary to a tectonic regime. the deposition of recognisable thrust-sheet tips in the piggyback basins supports the concept of a continuous thrust-fault process. proglacial and subglacial deformation glaciotectonic analyses distinguish between deformation generated in proglacial and in subglacial regimes (aber 1982; croot 1988; aber et al. 1989; pedersen 1993, 1996, 2000). it has already been argued that the thinskinned thrust-fault deformation of the rubjerg knude glaciotectonic complex is an example of proglacial deformation. the key evidence for this is the presence of intimately associated syntectonic piggyback basins. these basins must have been situated in front of the ice margin, with sedimentation taking place under open water, simultaneously with thrust-fault propagation. however, the subglacial deformation is represented locally by the 1 m thick glacitectonite occurring below the glaciotectonic unconformity that truncated the thrust-fault complex. this is found in the glaciolacustrine beds at the top of the ul02 thrust sheet in the northern part of the ulstrup section. it can be argued that here the subglacial deformation penetrated down a depth of c. 5 m below the glaciotectonic unconformity. mud diapirism and hydrodynamic brecciation occurred in this setting, probably caused by loading when the ice sheet overrode the sediments. the focus of subglacial deformation is at the blå-unconformity (fig. 32). in the northernmost 300 m of the cross-section, the effects of mud mobilisation increase to a point at which primary sedimentary as well as early structural features are completely destroyed. this phase of deformation is interpreted to have taken place while the sole of the frontal part of the ice sheet was fixed to the trailing end of the thrust-fault complex. this also implies that the velocity of the thrust faulting was equal to the advance of the ice sheet. the advance of the ice-sheet load corresponds to the mechanics of gravity spreading (pedersen 1987). the increasing propagating stress resulted in increasing mud mobilisation, and subsequently the overpressure was transmitted laterally by the mud fluid towards the foreland. the mechanism might well be compared to squeezing toothpaste out of its tube. at a certain stage, the fluid pressure was released, probably due to migration of all the hydrodynamic breccias, and the mobilised mud consolidated. subsequent to consolidation, the frontal sole of the icesheet released contact with the blå-unconformity and propagated over the thrust-fault complex formed in the foreland of the ice margin. during this process, subglacial shearing affected the top of the structureless consolidated mud, and anastomosing as well as plane-parallel shear fractures were formed (fig. 32). in soft sediment structural geology, the gravityspreading model has been successfully applied to the geological setting of the mud lumps in the mississippi delta (morgan et al. 1968; pedersen 1987; aber et al. 1989). it could therefore be suggested that a gravityspreading model due to clastic progradation might be the deformation mechanism. however, there is no known delta setting at this time/place that could have provided the basis for this model, and furthermore, the sand units observed here only reach a third of the thickness of the 100 m delta-sand units in the mississippi delta setting. finally, the presence of the glaciotectonic unconformity and related glacitectonite is incompatible with a sand sediment-spreading process. although a delta setting has not been documented, it might be suggested that a slope similar to that of a megascopic delta foreset existed, and that the deformation was caused by major gravity gliding on this slope, or was simply due to uplift in the hinterland. however, this is not considered likely. the isostatic rebound documented from the elevation of the vendsyssel formation reaches 60 m a.s.l. to this must be added the uplift due to the lowering of sea level; the area in the hinterland was thus an area of subsidence rather than uplift. structurally, a gravity-gliding model would provide extensional normal fault systems in the trailing end of the thrust-fault complex (pedersen 1987). this is not compatible with the observed increase in compressional structures in the hinterland, as documented in the cross-section and indicated by the balanced cross-section; a gravity-gliding model for the complex can therefore be rejected. such larger glaciotectonic complexes are often so impressive that some geologists suggest that they were formed by orogenic activity (lykke-andersen 1992; k. binzer, personal communication 1997). disregarding the obvious glacial geological indications, there are two features that distinguish glaciotectonic complexes from basement-involved deformation: (1) the superficial detachment, and (2) the rate of translation. in the rubjerg knude glaciotectonic complex, there are no infracrustal rocks involved and the thrust sheets are not rooted down into a deep-seated hinterland source. the lowermost detachment level is 40 m be183 low the reference level, which is more or less coincident with present sea level, and there are no indications that the deformation extended below the 40 m level. the velocity of thrust-sheet motion in orogenic mountain ranges is of the order of 1 cm per year (wiltscko & dorr 1983). in glaciotectonic systems, the velocity can be up to 100 times as fast, as documented by the velocity estimate of 2 m per year for the rubjerg knude glaciotectonic complex. glacial geological conditions the rubjerg knude glaciotectonic complex is interpreted to have formed due to the advance of the norwegian ice in the late middle weichselian. the norwegian ice melted back at the beginning of late weichselian time and was succeeded by a renewed advance of the scandinavian ice sheet from central sweden. in that part of denmark east and north of the main stationary line (figs 1, 12), the direction of this advance was towards the south-west and the advance is thus referred to as the ne-ice (houmark-nielsen 1987). the eastward advance of this ice towards vendsyssel probably formed the n–s-trending hilly landscape named jyske ås, the formation of which was contemporaneous with deposition of the outwash plain represented by the ribjerg formation. when the ne-ice advance reached the rubjerg knude glaciotectonic complex, it only resulted in minor superimposed deformation. the oblique orientation of the fold axis of the megaslump in the stenstue rende section might be due to such superimposed deformation, but in general very few glaciotectonic disturbances can be related to the ne-ice advance. that overriding by the ne-ice had so little ef fect may be attributed to strain hardening due to the preceding deformation, or the smoothing out of the landscape by the former glaciotectonic unconformity, which would facilitate the second overriding of the complex. ground-frozen conditions could also have been a factor, since this would have prevented drainage from the ice sheet through the substratum, resulting in high pore-water pressures at the sole of the ice. the effect of this would have been to facilitate easy and fast propagation over the complex, although it by then formed a hill in the landscape. the norwegian ice produced a hill-and-hole pair with rubjerg knude as the hill and the depression extending from lønstrup northwards as the hole. immediately after the melting back of the ne-ice, the landscape was covered by the vendsyssel formation. a contour map of the base of the vendsyssel formation (fig. 125) thus provides a picture of the geomorphology of the young glacial landscape unaffected by the succeeding 15 000 years of erosion. in fig. 125, the hill-and-hole pair is readily identified and the general e–w morphological trends are well represented. to the east, this trend is truncated by a strong se–nw hill-and-hole geomorphology, related to the ne-ice. the trend of the thrust-fault belt of the rubjerg knude glaciotectonic complex can be followed from the coastline to the east for about 2.5–5 km. the eastern fringe of the complex has been eroded down to sea level, probably by the ne-ice, and subsequently concealed by the vendsyssel formation. 184 185 facing page: fig. 125. contour map of the pre-vendsyssel formation landscape; for location, see fig. 13. note the depression north of lønstrup which represents the hole in the hill-and-hole pair morphology of a glaciotectonic complex; the corresponding hill is represented by the high at rubjerg knude. the map is based on data from the geus well database and from plate 1. conclusions structural analysis of the rubjerg knude glaciotectonic complex, based on detailed photogrammetric measurements and field investigations, provides a geological cross-section through a low-friction thrustfault system. interpretation of the entire thrust-fault architecture included unexposed parts of the complex, and is based on the construction of a balanced crosssection. a model for the dynamic development demonstrates that deformation progressed continuously and involved formation of duplexes and mud diapirs. although the thrust-fault structures were formed in a proglacial regime related to the advance of the norwegian ice (30 000 – 26 000 b.p.), the structures can be viewed as representing an almost complete model of thin-skinned thrust-fault systems. for descriptive purposes, the complex is subdivided into 13 sections, which demonstrate the structural development from a proximal to a distal position in the thrust-fault system. investigation of these sections provided the following main conclusions. 1. the structural elements in the rubjerg knude glaciotectonic complex comprise ramps and flats related to hanging-wall and footwall positions, respectively. hanging-wall anticlines and footwall synclines were formed due to thrust-fault propagation. back-thrust faults were formed during upper ramp-hinge propagation, and an irregular fold framework developed in relation to sequential duplex imbricate formation during footwall ramp collapse. foreland-dipping normal faults were formed in relation to translation of duplex segments. 2. from the balanced cross-section, the shortening during thrust-fault deformation is calculated to have been c. 50%. about 11% of the initial stratigraphic unit subjected to thrust faulting is estimated to have been lost due to erosion. the décollement zone was at its deepest position (40 m) in the proximal sections, becoming shallower towards the foreland. stacking of duplex segments is correlated with space problems created in the subsurface due to initial displacements at the upper levels. stacking of duplex segments correlates well with the elevation of the reference level in the system. 3. hydrodynamic brecciation was dominantly related to the hanging-wall ramps and flats. polydiapirism and mud mobilisation characterise the thrust zones. mud mobilisation resulted in the formation of larger mud diapirs, and preferentially evolved during hanging-wall propagation from the décollement level up above sets of intermediate and upper footwall ramps. 4. syntectonic deposition took place in piggyback basins overlying the thrust sheets. thrust sheets exposed to erosion provided sediment to the basins, and in some cases major lumps derived from the tips of thrust sheets slumped and slid as megablocks into the piggyback basins. 5. the thrust-fault deformation was caused by gravity spreading at the front of an advancing ice sheet. over-pressured mud formed an important part of the stress transfer. the average velocity of the thrust-fault displacement is estimated to have been 2 m per year. a 40 m thick succession of flat-lying sediments, extending for 12 km, was compressed into a thrust-sheet complex that was 6 km in length and up to 80 m thick. acknowledgements the geological survey of denmark and greenland is thanked for supporting this project during the last 10 years. initial investigations were carried out while the author held a senior stipend at the geological institute, university of copenhagen. the carlsberg foundation supported the project with a one year research grant, which is gratefully acknowledged; the danish research agency is thanked for financial support for printing this bulletin. keld dueholm and the institute of survey and photogrammetry are thanked for their co-operation and willingness to provide time and fa186 cilities at the photogrammetric instrument at the technical university of denmark. frants von platen-hallermund is thanked for assistance with the arc-info transformation and arc-view editing, which provided the graphic display of plates 1 and 2. alice rosenstand and benny m. schark helped in drafting the figures, and the staff of the graphic section at geus are thanked for technical support. i am indebted to a.k. higgins for reading the first draft of this manuscript, to two anonymous referees and the editor, jon r. ineson, for their constructive and helpful comments on the manuscript, and to my wife, gunver k. pedersen, for support and inspiring discussions during the progress of this work. references aaris-sørensen, k. 1995: palaeoecology of a late weichselian vertebrate fauna from nørre lyngby, denmark. boreas 24, 355–365. aaris-sørensen, k. & petersen, k.s. 1984: a late weichselian find of polar bear (ursus maritimus phipps) from denmark and reflections on the paleoenvironment. boreas 13, 29–33. aber, j.s. 1982: model for glaciotectonism. bulletin of the geological society of denmark 30, 79–90. aber, j.s. 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: a surface along which an overlying block is displaced relative to an underlying block. relative to bedding, two different elements are distinguished in a thrust fault: the ramp and the flat. ramp: a thrust-fault ramp cuts up-section in the direction of slip and dips towards the hinterland. the angle between bedding and the ramp is in general between 20° and 30° and will not exceed 45° due to general rules of initial fracturing. a ramp is linked to a lower flat at the lower ramp hinge and to an upper flat at the upper ramp hinge. a ramp may become foreland-dipping in special cases, mainly related to transport along an upper flat. flat: a thrust-fault flat is a bedding-parallel slip surface along which lateral displacement takes place. the lowermost thrust-fault zone in deformation complexes is in general referred to as the décollement surface, décollement zone or décollement level. in the present description the ‘décollement level’ is the term used for the thrust fault between a thrust sheet and an undisplaced footwall block below a footwall flat. as flats develop at dif ferent levels, the flats above the décollement level are referred to as intermediate flats and the upper flat (identical with the roof thrust fault). the thrust-fault flats are referred to by their depth from the upper reference zero-level indicated from the balanced section. thus the 20 m flat level is the horizontal thrust fault situated 20 m below the top reference level and 10 or 20 m above the décollement level. thrust sheet: a thrust sheet is the block displaced over a thrust fault. in this study, the thrust sheets are annotated according to the section in which they occur with two capital letters, referring to the named section, and a number referring to its position from leading edge to trailing end of the section. thus, kr01 is the thrust sheet nearest to the foreland in the kramrende section. a thrust fault is referred to according to the thrust sheet it displaces. a thrust sheet is synonymous with the hanging-wall block. hanging-wall block : the rock mass displaced over a thrust fault is a hanging-wall block. at the base, a hanging-wall flat and a hanging-wall ramp bound the hanging-wall block. at the roof, the hanging-wall block is capped by a top surface or a roof thrust fault. the roof thrust fault may constitute a footwall flat as well as a footwall ramp. footwall block: the rock below a thrust fault is a footwall block. the footwall block is bounded by a footwall ramp, and the top of the footwall block constitutes a top surface and/or a footwall flat. hanging-wall ramp: the segment of a ramp that bounds the hanging-wall block is a hanging-wall ramp. at the incipient displacement along a ramp, the hanging-wall ramp is thrust along a footwall ramp. when the hanging-wall ramp passes the upper ramp hinge, the hanging-wall ramp is thrust along a footwall flat. a hanging-wall anticline is always formed above a hanging-wall ramp. hanging-wall flat: the bedding-parallel thrust-fault boundary below the hanging-wall block is a hangingwall flat. when a hanging-wall flat is thrust up along a footwall ramp, the hanging-wall flat is re-orientated and becomes inclined towards the hinterland. when the hanging-wall flat is thrust along an upper footwall flat, the thrust fault again becomes bedding parallel. footwall ramp: the inclined thrust-fault boundary of a footwall block is a footwall ramp. the footwall ramp is either the ramp boundary to the undisplaced foreland or it forms the trailing ramp boundary of a thrust sheet. in this study, the trailing footwall ramp is referred to using the annotation of the thrust sheet/ footwall block that underlies it. thus the kr02 hanging-wall ramp is displaced up along the kr01 footwall ramp. footwall flat: a footwall flat is always the top of a footwall block. a footwall flat is more or less horizontal unless it is re-orientated during the displacement of a thrust sheet up along a ramp. hanging-wall anticline: when a hanging-wall block is thrust over an upper ramp hinge, the hanging-wall block is folded into an anticline with a foreland-dipping forelimb and a hinterland-dipping backlimb. this fold may also be termed a ramp anticline. during the progress of thrusting along the upper limb, the hanging-wall anticline develops into a flat-topped anticline. the flat-topped anticline may alternatively be regarded as a flat-lying thrust sheet with a foreland-dipping forelimb or a frontal thrust-sheet nose. however, it is important to note that above a hanging-wall ramp thrust along a footwall flat, a foreland-dipping surface is formed. 191 footwall syncline: when a thrust fault propagates up along a ramp, an anticline–syncline pair is formed above, and in front of, the tip of the thrust fault, identical to the formation of a fault-propagation fold. when the thrust fault finally breaks through the folded layers, the fold pair is separated into a hanging-wall anticline and a footwall syncline. a footwall syncline therefore represents the gentle deformation below the footwall ramp; this deformation does not add significantly to the displacement along the thrust fault. the footwall syncline may also be regarded as a drag fold. the case where this is the only correct interpretation is along a growth fault. here the sediments deposited syntectonically up against a hanging-wall ramp are successively bent into an overturned syncline. the footwall syncline is identical to a trailing syncline. duplex: a duplex is one or more thrust-sheet segments entirely bounded by thrust faults and thus overlain by a thrust sheet. a thrust sheet bounded by thrust faults is called a horse, originally regarded as a minor rootless thrust-sheet segment. some of the lower thrustsheet segments described in this study are identical to horses, although the more neutral term ‘segment’ is adopted here. the formation of a duplex is related to the ‘footwall ramp collapse’ (boyer & elliott 1982), whereby progressive failure during thrust-fault propagation creates successively younger thrust faults below older ones. a parcel of thrust-sheet segments may be stacked to form a duplex complex. imbricate fan : a branching thrust-fault complex in which the individual thrust faults reach the surfaces or top level is called an imbricate fan. an imbricate fan is termed a duplex if the upper boundary is a roof thrust. antiformal stack: when a duplex is fault-bend-folded over a footwall ramp, an antiformal structure similar to a hanging-wall anticline is formed. due to the complex stratigraphic relationship within such a structure, it is referred to as an antiformal stack. piggyback thrusting: when an older thrust sheet rests on the back of a younger thrust sheet and is transported due to the displacement along the thrust faults bounding the younger thrust sheets, it is called piggyback thrusting. piggyback basin: just as piggyback thrusting refers to transport of a thrust sheet, the term is also applied to a basin that accumulates sediments during translation on the back of an active thrust sheet: the piggyback basin (ori & friend 1984; ricci lucchi 1986). in this study, the term is mainly used in the description of an area of sedimentation between two thrust sheets. in general, the piggyback basin is deposited between a fault-bend thrust-sheet tip in the distal part of a thrust structure and bounded by a hanging-wall ramp at the proximal boundary of the basin. the term piggyback basin can only be applied to successions identified as having been deposited syntectonically. 192 appendix 2 specification of photogrammetric work the construction of the rubjerg knude cross-section (plate 1) is based on a multi-model photogrammetric investigation of the cliff section, with the application of the method described by dueholm (1992). a series of oblique photographs were taken from a cessna fixed-wing aircraft in june 1993. the camera used for the photography was a minolta xg2, which had been tested and calibrated for its optical specifications at the laboratory of photogrammetry at the danish technical university. the films used were standard 24 × 36 mm colour diapositive. the photographs were taken with 66% overlap from a distance of 200–300 m with an inclination angle of c. 35°. from the series of photographs, 70 samples were selected for setting up three sets of templates, which included 67 stereoscopic models. in the laboratory, the orientation of the stereo-models was carried out based on ground control points adapted from two sets of vertical aerial photographs at a scale of 1:25 000, namely d9202 g 1365–66 and kms 9203 a509–10 taken in may 1992. the strike of the section line is n15°e from rubjerg knude and southwards. north of rubjerg knude, the strike is n24°e, which is nearly parallel to the direction of the coastline along the beach. fortunately, this is also a reasonable approximation of being perpendicular to the main concentration of structural strikes (bedding, thrust faults and fold axes; fig. 10). a minor adjustment of the northern and southern section lines was subsequently implemented to make the cross-section fit to the general plane of orthographic projection with a projection axis striking 107°. the stereoscopic instrument used for the investigation was a kern dsr 15 analytical plotter with a dec vms operating system and the special attached geoprogram developed by dueholm (1992). five different labels were used for the features outlined by the floating mark: line type 1 includes bedding traces, line type 2 includes the main unconformities, line type 3 was used for the contacts between geological units (members and formations), line type 4 outlines thrust faults, and finally line type 5 was used for topographic features (dunes, scree cones, strandplain, rockfalls etc.). digitalisation of the geological structures in the stereo-models was administrated in data files, each covering a plot-area. the plot-areas covered 500 m of the rubjerg knude cross-section, and 13 plot-areas were used for the analogue plotting of data digitised in the stereo-model. the digital data were stored for the later construction of the cross-section and the transformation for other programs applied for the management of the cross-section display. the orientation of models and setting up the system for the cross-section investigation took about one week, and the photo-geological compilation work was made over a period of three months in the autumn of 1993. the average progress was two models per day. the benefit of the multi-model analytical stereo-plotter is that features can be traced continuously from one model to the adjacent models. thus one is not restricted to working model by model, but the compilation can be extended over several models using the same set of templates. by january 1994, the cross-section could be plotted out in a normal vertical projection profile plan from the stored digital data with the application of the program facilities prepared by dueholm (1992). the scale of the rubjerg knude crosssection in the draft versions is 1:500, and the accuracy of the plotted data is estimated to be better than 25 cm. in 1995–1996, the cross-section details observed in the photo-geological models were checked in the field, and in 1997 the templates were set up again for correction, adjusting and compilation of details in the cross-section. geological survey of denmark and greenland bulletin 4, 2003, pp 81-84 81 the climate of europe is strongly influenced by heat transport by ocean currents flowing from equatorial regions towards the arctic (clark et al. 2002). during recent years, research has been increasingly focused on factors affecting this circulation, e.g. the freshwater budget of the arctic which is influenced by glacial meltwater from north and east greenland outlet glaciers (linthout et al. 2000, mayer et al. 2000). furthermore, the climate is affected by snow cover that, apart from its contribution to the freshwater budget, provides feedback effects in that it reflects most of the solar radiation. apart from arctic sea-ice cover, the greenland ice sheet is the largest permanent iceand snow-covered area in the northern hemisphere, with an area of 1.67 × 106 km2 and by far the largest storage of ice with a volume of 2.93 × 106 km3 (bamber et al. 2001). most of the mass loss from the greenland ice sheet (the least known mass-balance parameter) occurs in the marginal region of the ice sheet, which is also the area where the largest changes in albedo occur. the geological survey of denmark and greenland (geus) has for many years carried out research along the greenland ice sheet margin to monitor changes of mass balance and melt conditions. changes in mass balance: present knowledge changes in mass balance over the interior part of the greenland ice sheet can conveniently be assessed from snow pits and ice cores, because the ‘history’ is preserved as annual layers in the accumulation zone. at the ice margin this stratigraphic approach cannot be used due to melting of the surface. monitoring of ablation is therefore carried out using stakes drilled into the ice. until about 1980, mass balance measurements in greenland were few and of short duration. during the 1980s, intensive studies were initiated, motivated mainly by hydropower investigations (weidick 1995). this was followed by a period in the early and middle 1990s when climate-related mass-balance campaigns were carried out in north-east and north greenland in conjunction with geological mapping expeditions (bøggild et al. 1994; thomsen et al. 1997). an overview of the various field activities is given in fig. 1. an analysis of field results together with model estimates has yielded ablation rates varying from up to 10 m/yr in the southernmost part of greenland to a few metres per year in north-east greenland (huybrechts et al. 1991). since the traverse of the ice sheet during the expédition glaciologique internationale au groenland (e.g.i.g.) in 1957–1960, locations on this traverse route have been used for several investigations, including long-term surface-change and mass-balance studies (finsterwalder 1959; fischer et al. geological survey of denmark and greenland bulletin 4, 81–84 (2004) © geus, 2004 towards an assessment of the balance state of the greenland ice sheet carl e. bøggild, christoph mayer, steffen podlech, andrea taurisano and søren nielsen fig. 1. locations on the greenland ice sheet margin where mass balance investigations have been carried out, including recent activities by the authors. 1995). major advances in the assessment of the greenland ice sheet mass balance were achieved during the program for arctic regional climate assessment (parca) between 1991 and 1999. the project was initiated with a balance assessment around the 2000 m altitude contour on the ice sheet, and was gradually expanded with add-on projects into a large-scale, multidisciplinary project to determine the balance state of the ice sheet (thomas et al. 2001). these largescale investigations have incorporated both visible and radar satellite images, airborne radar depth sounding, multi-parameter core analysis, weather stations, climate-modelling etc. most studies have been carried out in the interior parts of the ice sheet, where spatial variability of accumulation, ice thickness and other factors are small compared to the marginal regions. hence, coverage over large areas can more easily be obtained from remote sensing data, combined with a few observations and ground-truth calibrations. the parca campaign strongly improved the knowledge of snow accumulation distribution and provided a clear estimate of ice thickness and volume of the ice sheet. although rather large spatial and temporal variability in local accumulation was demonstrated, the interior of the ice sheet as a whole was shown to be in approximate balance (thomas et al. 2001). however, airborne laser altimetry extended observations further towards the ice margin and documented relatively rapid thinning particularly in the south to south-eastern and north-western parts of the ice sheet (fig. 2; krabill et al. 1999). in the late 1990s, an analysis of elevation change history and mass balance was carried out by geus at two confined outlet glaciers, sermilik bræ in south greenland and qamanaarsuup sermia in west greenland (figs 1, 3). both studies confirm the observed thinning from laser altimetry by krabill et al. (1999). the laser altimetry results are, however, restricted to the period from 1991 to 2001, whereas the geus elevation studies could be extended back to around 1950 using topographic models based on the earliest aerial photographs available. this longer time frame of the geus studies provides a more reliable estimate of the balance state, since the effects of year-to-year variability are reduced. preliminary results show that the average thinning of the ice sheet margin over the last five decades may locally exceed the laser altimetry estimates of 1000 mm/yr in the areas of investigation. the observed thinning at sermilik bræ ranges from 2000 to 8000 mm/yr, while at qamanaarsuup sermia thinning rates of around 1000 mm/yr have been found. the thinning appears to be only in part due to increased melting at both sites. towards automatic ablation observations the observed strong thinning of the greenland ice sheet margin has called for the development of new automatic data collection methods. traditionally, these measurements have been made with stakes drilled into the ice that were periodically visited to determine surface lowering by melting. however, the stakes need to be re-drilled at least once during the ablation season in order to secure a continuous time series. this is a highly labour-intensive and costly way to retrieve ablation data, particularly in view of the approximately 5000 km long perimeter of the ice sheet margin. to resolve this dilemma an automatic mass balance station (ams) has been under development at geus since 2001 82 fig. 2. elevation changes in cm/yr on the greenland ice sheet from airborne laser altimetry detected during the parca campaign (ref.: http://aol.wff.nasa.gov/aoltm.html). blue colour tones indicate areas of thinning. (mayer et al. 2002), which, with a new concept of ablation measurements, will allow uninterrupted operation for several years. the most recent design combines a sturdy tripod with a specially developed sensor system that is coupled to an automatic, satellite-based data transmission system enabling data retrieval at six-hour intervals (bøggild et al. in press). future perspectives present knowledge of the balance state of the ice sheet reveals that the ice margin is thinning rather rapidly in the southeastern and north-western parts of the ice sheet (krabill et al. 1999). this is in contrast to earlier sensitivity estimates of ice sheet thinning during a warming climate (huybrechts et al. 1991), that indicated the south-western and north-eastern marginal regions to react most strongly by thinning of the margin. our present understanding of the climate sensitivity of the ice sheet margin is evidently insufficient. surface observations are unfortunately lacking from several areas showing the most pronounced thinning in remote sensing data, including melville bugten in north-west greenland and the ice margin south of tasiilaq, east greenland (fig. 2). studies of the cryosphere (snow and ice on sea and land) have gained considerable focus in recent years, and new satellites have been developed specifically dedicated to monitor elevation changes from space. in 2003, nasa (national aeronautics and space administration) launched the icesat satellite, and in 2005 esa (european space agency) will follow with the cryosat satellite. however, the satellites only detect elevation changes and not the underlying causes. additional in situ observations of climate and mass balance using ams are needed for determining the annual mass balance ‘turn-over’ and detect possible changes. the new stations will be able to measure ice movement as well as ablation, although observations of elevation changes over several years will be required in order to derive long-term trends not biased by inter-annual fluctuations. alternatively it may be possible to develop digital elevation models by making use of the danish archives of aerial photographs dating back to about 1950. with the launch of new satellites and a more widespread use of ams a detailed monitoring of the cryosphere becomes more realistic. recently, a european union programme (euroclim) was initiated to develop a climate change monitoring system focusing on the cryosphere. one part of this system (under development at geus) aims at detecting changes in surface conditions of the greenland ice sheet (fig. 4). furthermore, initiatives have been taken towards integrating landand sea-ice studies to determine the climate sensitivity of the cryosphere and, not least, the impact of the cryosphere on the climate system. both marine and terrestrial ice show similar interactions with climate. 83 fig. 3. observed changes at sermilik bræ and qamanaarsuup sermia in historical time. a: fjord and glacier front of sermilik bræ in 2001 recorded by aster imagery. the lateral margins of the drainage area are indicated by thin black lines. b: terminus changes of qamanaarsuup sermia since the last advance during the little ice age (1880–1890; beschel 1961). the vertical aerial photograph was taken by the danish geodetic institute in 1968 (reproduced by permission a.200/87). 84 references bamber, j.l., layberry, r. & gogineni, s.p. 2001: a new ice thickness and bed data set for the greenland ice sheet. journal of geophysical research 106(d24), 33773–33780. beschel, r. 1961: dating rock surfaces by lichen growth and its application to glaciology and physiography (lichenometry). in: raasch, g. (ed.): geology of the arctic, 1044–1062. toronto: university of toronto press. bøggild, c.e., reeh, n. & oerter, h. 1994: modelling ablation and massbalance sensitivity to climate change of storstrømmen, northeast greenland. global and planetary change 9, 79–90. bøggild, c.e., olesen, o.b., ahlstrøm, a.p. & jørgensen p. in press: automatic glacier mass balance observations using pressure sensors. journal of glaciology. clark, p.u., pisias, n.g., stocker, t.f. & weaver, a.j. 2002: the role of the thermohaline circulation in abrupt climate change. nature 415, 863–869. finsterwalder, r. 1959: expédition glaciologique internationale au groenland 1957–1960 (e.g.i.g.). journal of glaciology 3(26), 542–546. fischer, h., wagenbach, d., laternser, m. & haeberli, w. 1995: glaciometeorological and isotopical studies along the egig-line, central greenland. journal of glaciology 41(139), 515–527. huybrechts, p., letreguilly, a. & reeh, n. 1991: the greenland ice sheet and greenhouse warming. global and planetary change 3(4), 399–412. krabill, w., frederick, e., manizade, s., martin, c., sonntag, j., swift, r., thomas, r., wright, w. & yungel, j. 1999: rapid thinning of parts of the southern greenland ice sheet. science 283, 1522–1524. linthout, k., troelstra, s.r. & kuijpers, a. 2000: provenance of coarse icerafted detritus near the se greenland margin. netherlands journal of geosciences / geologie en mijnbouw 79(1), 109–121. mayer, c., reeh, n., jung-rothenhäusler, f., huybrechts, p. & oerter, h. 2000: the subglacial cavity and implied dynamics under nioghalvfjerdsfjorden glacier, ne-greenland. geophysical research letters 27(15), 2289–2292. mayer, c., bøggild, c.e., podlech, s., olesen, o.b., ahlstrøm, a.p. & krabill, w. 2002: glaciological investigations on ice-sheet response in south greenland. geology of greenland survey bulletin 191, 150–156. thomas, r.h. & parca investigators 2001: program for arctic regional climate assessment (parca): goals, key findings, and future directions. journal of geophysical research 106(d24), 33691–33706. thomsen, h.h., reeh, n., olesen, o.b., bøggild, c.e., starzer, w., weidick, a. & higgins, a.k. 1997: the nioghalvfjerdsfjorden glacier project, north-east greenland: a study of ice sheet response to climate change. geology of greenland survey bulletin 176, 95–103. weidick, a. 1995: greenland. in: williams, r.s.jr. & ferrigno, j.g. (eds): satellite image atlas of glaciers of the world. u.s. geological survey professional paper 1386-c, 141 pp. authors’ addresses c.e.b., c.m., s.p. & s.n., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ceb@geus.dk a.t., norwegian polar institute, polarmiliøsenteret, n-9296 tromsø, norway. fig. 4. surface conditions over the greenland ice sheet from modis imagery from august 2002, as part of the euroclim project. the surface is divided into several classes, ranging from pure ice at the ice sheet margin to dry snow in the interior. calibration of the results requires confirmation by ground observations and high-resolution satellite data. << /ascii85encodepages false /allowtransparency false /autopositionepsfiles true /autorotatepages /all /binding /left /calgrayprofile (dot gain 20%) /calrgbprofile (srgb iec61966-2.1) /calcmykprofile (u.s. web coated \050swop\051 v2) /srgbprofile (srgb iec61966-2.1) /cannotembedfontpolicy 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<feff4f7f752890194e9b8a2d5b9a5efa7acb76840020005000440046002065874ef65305542b8f039ad876845f7150cf89e367905ea6ff0c4fbf65bc63d066075217537054c18cea3002005000440046002065874ef653ef4ee54f7f75280020004100630072006f0062006100740020548c002000520065006100640065007200200035002e0030002053ca66f465b07248672c4f86958b555f3002> >> >> setdistillerparams << /hwresolution [2400 2400] /pagesize [595.000 842.000] >> setpagedevice geological survey of denmark and greenland bulletin 33, 2015, 45-48 45 magma mixing, mingling and hybridisation at different crustal levels: snapshots from 1.9 billion years of magmatism in south-eastern greenland thomas f. kokfelt, samuel m. weatherley, jakob k. keiding and trygvi b. árting during fi eld work in 2014, we investigated a suite of igneous intrusions in south-eastern greenland between 65° and 67°n. many of the intrusions show widespread evidence for juxtaposition of diff erent magmas in the liquid state and subsequent mixing, mingling and hybridisation. here we present fi eld evidence for these processes from three areas that diff er in age and geological setting. we discuss the signifi cance of mingling, mixing and hybridisation features in the fi eld area, motivated by their abundance in the area, the morphological variation between intrusions that were emplaced at diff erent crustal levels, the implications for magma genesis in collisional and rift settings, and the implications for the interior dynamics of igneous bodies. magmatism in south-eastern greenland th e magmatic history of south-eastern greenland broadly falls into two episodes: (1) the proterozoic nagssugtoqidian orogeny and preceding subduction, and (2) the palaeogene opening of the north atlantic. th e nagssugtoqidian orogen formed by a ne–sw collision between the rae and the north atlantic cratons that started at c. 1880 ma and ended with a post-orogenic collapse at c. 1740–1690 ma (see kolb 2014 and references therein). several intrusions and intrusive complexes are associated with the orogeny (fig. 1). th ese include the c. 1885 ma ammassalik intrusive complex, interpreted as originating in an island arc, several post-tectonic granites, and the intermediate to felsic ikaasartivaq intrusive complex with an uncertain intrusion age of 1680 ma (kalsbeek et al. 1993). th ree separate intrusions of diorite and tonalite are found north of ikaasartivaq strait. th e ages of these palaeoproterozoic intrusions, their geotectonic setting and relationship to the other intrusive bodies are poorly constrained. intrusive complexes from the palaeogene extend from 66° to 75°n. th ey were emplaced between c. 55 and 25 ma (larsen et al. 2014 and references therein) and encompass mafi c to felsic compositions of alkaline to tholeiitic affi nity in addition to carbonate-related lithologies (nielsen 2002). th e intrusive complexes and associated coeval lavas and coastparallel dyke swarms are all part of the palaeogene north atlantic igneous province (naip) that formed as a result of oceanic rift ing under the infl uence of the ancestral icelandic mantle plume (brooks 2011). in this contribution, we present evidence for magma mingling, mixing and hybridisation in south-eastern greenland, focusing in particular on the ammassalik intrusive complex, the ikaasartivaq intrusive complex and the kialineq intrusive centre. magma mixing, mingling and hybridisation magma mixing is a recognised process operating in most magmatic systems worldwide and may be a trigger for generating volcanic eruptions (sparks & sigurdsson 1977). th e term ‘mixing’ refers to a process where two magmas blend to form a new magma of intermediate composition. th e extent to which the magmas blend or mix critically depends on diff erences in the rheological properties and the relative proportions of the juxtaposed magmas (sparks & marshall 1986). important parameters that determine the rheology of a given magma include temperature, composition and crystal load. for large property contrasts, the mixing process will © 2015 geus. geological survey of denmark and greenland bulletin 33, 45–48. open access: www.geus.dk/publications/bull kialineq intrusive centre tasiilaq intrusion ikaasartivaq intrusive complex ammasalik intrusive complex tasiilaq 67°n 34°w laube gletscher laube gletscher syenite intrusion ikaasartivaq strait palaeogene basalt 50 km greenland pre-cretaceous rocks cretaceous/palaeogene sedimentary rocks intrusion desribed here other intrusions fig. 1. geological map of south-eastern greenland showing some of the prominent intrusions and intrusive centres in the region. the intrusions described here are the c. 1885 ma tasiilaq intrusion of the ammassalik intrusive complex, the c. 1680 ma ikaasartivaq intrusive complex and the c. 35 ma kialineq intrusive centre. +10 –8 4646 typically be incomplete and the original magmas remain as identifi able units; the result is a mingled magma. hybridisation refers to the general mixing process that may produce a range of variably heterogeneous intermediate magmas. classical examples of magma mixing and mingling include eroded plutonic systems within the naip including iceland (brooks 2011). a common feature in many of these exhumed magma reservoirs is the occurrence of ‘net-veined’ complexes consisting of felsic, mafi c and hybrid (intermediate) rocks. tasiilaq intrusion of the ammassalik intrusive complex – th e tasiilaq intrusion, dated to 1886 ± 2 ma (hansen & kalsbeek 1989), is the central part of the palaeoproterozoic ammassalik intrusive complex that forms a c. 85 × 20 km wnw–ese-trending array of three large mafi c intrusions (fig. 1). th e intrusion mainly comprises leuconorite that cross-cuts melagabbro, and late anorthosite and hypersthene veins that cross-cut the leuconorite and melagabbro. th e wall rock is quartzofeldspathic garnet-rich gneiss, which has granulite facies assemblages close to the contact, and has been mobilised and mixed with the intrusive rock forming a marginal zone of hybrid rocks. based on thermobarymetry and fl uid inclusion data, andersen et al. (1989) concluded that the intrusion crystallised at c. 1000–1100°c and 6–8 kbar (c. 18–24 km depth), i.e. in the middle to deep crust. in the fi eld, melagabbro oft en forms 20–100 m wide trains within larger domains of leuconorite, defi ning the same general wnw–ese-orientation as the overall structures of the ammassalik intrusive complex. in some cases leuconorite has intruded melagabbro forming sharp angular contacts (fig. 2a); in other outcrops, the two lithologies form lenticular intermingled bodies with smeared out, wavy contacts and with no chilled margins (fig. 2b). ikaasartivaq intrusive complex – th e ikaasartivaq intrusive complex is composed of mafi c and felsic intrusions, which straddles the ikaasartivaq fj ord c. 25 km north of tasiilaq (fig. 1). th e complex measures c. 32 × 25 km and comprises rocks of granitic, dioritic and gabbroic composition. it was fi rst mapped by wright et al. (1973), who described it as a lateor post-orogenic calc-alkaline suite related to the uplift of the area. th ese authors proposed that the complex was emplaced into relatively shallow crustal levels based on observations of sharp contacts between intrusive units and negligible evidence for contact or retrogressive metamorphism. mixing and hybridisation features within the intrusive complex are ubiquitous and show a range of styles and morphologies. in the main body of the complex, mingling features between diorite and granite are observed at millimetreto metre-scales in the form of blobs and globules and occasionally as angular net-veined areas. although mingling features are quite pervasively developed throughout the complex, they tend to be more prominent and abundant in the western part of the intrusion. here they oft en occur as rather extensive pilfig. 2. examples of composite magmatic systems intruded at different crustal levels. a, b: the tasiilaq intrusion representing deep to middle crustal levels. a: leuconorite intruding melagabbro in a semi-molten state. b: mingling structures with lensoidal wavy contacts between leuconorite (no) and melagabbro (gb). c, d: the ikaasartivaq intrusive complex intruded at intermediate to upper crustal levels. c: a c. 5 m wide composite mafic-felsic dyke intruded into gneiss basement. d: contact of dyke with gneiss (gn) wall rock, marginal zone of hybridised magma (hyb), and interior part of dark, distorded diorite pillows (di) in pale granite (gr) matrix; note crenulated, chilled margins in diorite. e, f: the kialineq intrusive centre that represents a shallow intrusion (<5 km). e: a 4–5 m wide composite sheet of syenite with up to metre-sized mafic pillows in diorite. f: composite dyke in the syenite body (sy) at laube gletscher showing evidence of mingling and hybridisation on a small scale. di di a b c d e f gb no hyb gr gn sy hyb 47 low complexes that texturally resemble the mafi c-felsic mixed rocks in the kialineq intrusive centre (see below). th e complex also hosts a range of composite dykes and sills, some of which are texturally similar to the classical net-veined complexes in iceland. mingling and mixing features within these smaller bodies include caulifl ower textures, irregular margins between mafi c globules and felsic host melt (fig. 2c, d), and decimetreto several metre-scale mafi c pillows within felsic material. th e margins of intrusive features and edges of individual mafi c globules commonly show signs of hybridisation as indicated by intermediate coloured rocks (fig. 2d). kialineq intrusive centre – th e kialineq intrusive centre is located c. 200 km north of tasiilaq (fig. 1) in an extremely rugged terrain. it is dominated by gabbro-diorite-syenitegranite plutons and is cut by numerous coast-parallel mafi c dykes. many of the plutons are closely associated with extensive mafi c-felsic mixed magma complexes (brooks 2011 and references therein). only a few radiometric ages are available, but they appear to show a bimodal distribution with mafi c magmatism at 56–49 ma, represented by the mafi c imilik intrusion, followed by a hiatus and subsequent activity at 37– 35 ma of plutons with more evolved compositions (larsen et al. 2014). although the chronology is poorly constrained, it agrees with the fi eld observations and suggests a general sequence of emplacement of gabbro, mafi c dykes, mafi c-felsic complexes, syenite and granite. mafi c-felsic complexes, oft en referred to as net-veined complexes, are widespread and ubiquitous in the kialineq district, and have also been reported elsewhere from the naip (nielsen 2002). th e vertical extent of the mafi c-felsic complexes is rarely well exposed in the kialineq intrusive centre, but fi eld evidence suggests that these complexes can be at least 400 m thick. although the fi eld relationships are complex and vary substantially with locality, they show that the mafi cfelsic complexes are dominated by dioritic pillows surrounded by comagmatic syenites (fig. 2e). th e pillows range in size from c. 0.1 to 2 m in diameter and consist of fi neto mediumgrained diorite, typically coarsening from rim to centre with some pillows having chilled rims against the syenites. th is implies a signifi cant thermal gradient existed between the two liquids. more typical sensu stricto net-veined textures of angular mafi c blocks separated by rather homogeneous syenite matrix also occur. such breccia-type textures testify that the basic material in some cases was partly solidifi ed when the syenite was emplaced. in contrast, the pillow complexes suggest synmagmatic interaction and mingling that demonstrate the co-existence of mafi c and felsic liquids. hybridisation was observed at the laube gletscher (fig. 2f), but is generally rare in the area. interestingly, vesicles and centimetre-sized miarolitic cavities are common features in kialineq intrusive rocks, showing that exsolution of gaseous phases occurred. th ese exsolution features, along with the occurrences of ring dykes and bell-jar plutons, indicate a shallow emplacement depth in a subvolcanic (cauldron) environment. discussion, summary and outlook we have described three examples of mafi c-felsic complexes in south-eastern greenland that show distinct mingling features irrespective of diverse geological settings, level of emplacement in the crust and compositional range. here we briefl y discuss the main diff erences between the three areas in order to constrain the key responsible processes that generated the observed features. th e tasiilaq intrusion represents a deep crustal end-member of the study area. compared to the other two examples, the mingled lithologies of the tasiilaq intrusion (leuconorite and melagabbro) contrast less strongly in composition, and so large diff erences in solidus temperatures are unlikely. further support for this hypothesis is off ered by the lack of chilled margins at lithological contacts, although this could in part also refl ect the slow cooling of the intrusion. th e textural relationships between the two main lithologies are consistent with leuconorite being intruded into melagabbro at a point where the latter was partly solidifi ed. th e mingling-like features are oft en smeared out defi ning a general wnw–ese orientation, suggesting a syn-magmatic infl uence from the regional stress fi eld of the orogen. overall, the tasiilaq intrusion seems to record a markedly diff erent and more tectonised image of mingling relations between magmas of minor compositional contrast than shown by the other two examples. th e mingling features found in the ikaasartivaq intrusive complex and the kialineq intrusive centre show many overall similarities, but also some notable diff erences. both areas represent fairly high crustal levels as opposed to the tasiilaq intrusion and contain an apparent bimodality of magma compositions (oft en referred to as the ‘daly gap’), including diorite and syenite/granite as end-members. th e common observation in both systems are chilled mafi c pillows or blobs in contact with felsic melts, indicating that a considerable temperature contrast existed between the respective mingling liquids. th is excludes a model in which the mingling components are formed by silicate liquid immiscibility processes that otherwise are believed to be important in many magmatic systems. only limited evidence for hybridisation is seen in the kialineq intrusive centre as compared to the common occurrence in the ikaasartivaq intrusive complex (fig. 2c vs. fig. 2e). th is could refl ect diff erences in a range of parameters, 4848 especially in the crustal level. since temperature gradients between magma and host rock are greater at shallow crustal levels, the time window for mixing and hybridisation in the kialineq system is expected to be shorter. th is may account for better preservation of end-member compositions in the kialineq intrusive centre than the other two examples studied. in all three areas studied, the scales at which mafi c and felsic components mingle to form blobs of diff erent sizes probably refl ect multiple factors, such as contrasts in viscosity, temperature and density of the mingling magmas, the relative abundance of mafi c to felsic magmas. it is beyond the scope of this paper to discuss these parameters in detail, but we suggest that blob development indicates an environment where diff erences in magma temperatures are small. unsurprisingly, these features are found in the largest intrusions. caulifl ower textures with grain size reduction indicate environments with stronger thermal gradients and these are typically found in the shallower environments or in smaller, sill or sheet-like bodies. in summary, igneous rocks in south-eastern greenland provide abundant evidence for bimodal magmatism, and magma mingling, mixing and hybridisation in diff erent tectonic settings and crustal levels. key diff erences between the three areas relate partly to various extents of mingling vs. hybridisation and partly to textural and structural characteristics of the interface between mafi c and felsic components (e.g. blob vs. pillow size, extent of grain-size reduction, degree of surface crenulation). th e governing factors are several, including (1) depth of emplacement which determines thermal conditions and cooling rates, (2) diff erences in magma-intensive parameters for the sets of dual end-members (temperature, density, viscosity, chemical composition, crystal contents), (3) the relative proportions of the mafi c and silicic end-members, and (4) size and geometry of the intrusive bodies (small dykes or sheets vs. large magma chamber systems). future work aims to qualify the reasons for the diff erences in more detail, to assess how the region fi ts into a broader understanding of magmatism and petrogenesis in diff erent tectonic settings and to investigate the coupled chemical and physical dynamics of juxtaposed magmas. one aspect will be to understand how the mingling features relate to the geological setting and intrusion shape, and investigate how these rocks compare to net-veined complexes in rift settings and andesites in arc and collisional settings. another aspect regards the origin of the apparent bimodality and understanding the role of magma mixing as a process masking the compositional gap. detailed petrological, geochemical and isotopic studies will be carried out to characterise endmember components in order to address the petrogenetic link between the mafi c and felsic components. th is will help to distinguish between models of fractional crystallisation versus partial melting of the archaean crust and to explain the apparent bimodality. acknowledgements th e work is part of the segment project that aims to evaluate the economic potential of the larger tasiilaq area and to gather geological information in general and is fi nanced by the ministry of industry and mineral resources in greenland and the geological survey of denmark and greenland. special thanks to christian tegner, chip lesher and th omas ulrich for collaborative fi eldwork in the kialineq area. references andersen, t., austrheim, h. & bridgwater, d. 1989: p–t and fl uid evolution of the angmagssalik “charnockite” complex, se greenland. in: bridgwater, d. (ed.): fluid movements – element transport and the composition of the deep crust, 71–94. dordrecht: kluwer academic publishers. brooks, c.k. 2011: th e east greenland rift ed volcanic margin. geological survey of denmark and greenland bulletin 24, 96 pp. hansen, b. & kalsbeek, f. 1989: precise age for the ammassalik intrusive complex. rapport grønlands geologiske undersøgelser 146, 46–47. kalsbeek, f., austrheim, h., bridgwater, d., hansen, b.t., pedersen, s. & taylor, p.n. 1993: geochronology of archaean and proterozoic events in the ammassalik area, south-east greenland, and comparisons with the lewisian of scotland and the nagssugtoqidian of west greenland. precambrian research 62, 239–270. kolb, j. 2014: structure of the palaeoproterozoic nagssugtoqidian orogen, south-east greenland: model for the tectonic evolution. precambrian research 255, 809–822. larsen, l.m., pedersen, a.k., tegner, c. & duncan, r.a. 2014: eocene to miocene igneous activity in ne greenland: northward younging of magmatism along the east greenland margin. journal of the geological society 171, 539–553. nielsen, t.f.d. 2002: palaeogene intrusions and magmatic complexes in east greenland, 66 to 75° n. danmarks og grønlands geologiske undersøgelse rapport 2002/113, 249 pp. sparks, s.r.j., sigurdsson, h. & wilson, l. 1977: magma mixing: a mechanism for triggering acid explosive eruptions. nature 267, 315–318. sparks, r.s.j. & marshall, l.a. 1986: th ermal and mechanical constraints on mixing between mafi c and silicic magmas. journal of volcanology and geothermal research, 29, 99–124. wright, a.w., tarney, j., palmer, k.f., moorlock, b.s.p. & skinner, a.c. 1973: th e geology of the angmassalik area, east greenland and possible relationships with the lewisian of scotland. in: park r.g. & tarney, j. (eds): th e early precambrian of scotland and related rocks of greenland, 157–177. keele: university of keele. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: tfk@geus.dk geological survey of denmark and greenland bulletin 35, 2016, 83-86 83© 2016 geus. geological survey of denmark and greenland bulletin 35, 83–86. open access: www.geus.dk/publications/bull the geological survey of denmark and greenland has carried out a detailed mapping project in the eastern baffin bay region covering c. 200 000 km2 (fig. 1). the purpose of the study was to update the previous mapping by using the most recent data and provide an improved basis for evaluation of the geological development and hydrocarbon potential of the region. after licensing rounds in 2007–2008 and 2010 considerable new geophysical and geological data were acquired in the region, and the extensive database now includes more than 100 000 km 2d seismic data and a number of wells (fig. 1). the results of the work are summarised below and suggest deep basins and large ridges with complex structures. geological setting in areas along the west greenland continental margin, a number of basins with proterozoic, cretaceous and cenozoic sedimentary successions have been identified (e.g., dawes 1997; dam et al. 2009; rolle 1985). a number of rifted basins with large structural highs are interpreted to have developed in the east baffin bay region (whittaker et al. 1997; gregersen et al. 2013). during the paleocene and eocene, oceanic crust developed in central baffin bay and the cretaceous rifted continental margin of west greenland was separated from eastern canada (oakey & chalmers 2012). the large-scale movements between greenland and canada generated new structures during the palaeogene and reactivated faults within cretaceous basins. results the study included seven sub-projects: (1) seismic interpretation and mapping; (2) well correlation; (3) biostratigraphy; (4) seismic facies analysis and seismic inversion; (5) overview of source rocks; (6) maturity modelling; and (7) structural development. this paper describes some key results from the seismic interpretation and structural development. interpretation of horizons and units were carried out with schlumberger petrel© software and included data from wells, seismic surveys, gravity surveys and magnetic surveys. additional data from seabed sampling and outcrops in the region were used for geological interpretation and as analogues. a robust framework was established with fourteen seismic stratigraphic horizons (a1–hx) and eleven mega-units (a–h). seismic cross-sections show deep, rifted basins separated by large structural highs (fig. 2). the lateral extends and topography of the basins and structures are outlined in depth-structure maps (fig. 3). in the kap york basin (fig. 4), the upper part of megaunit h can be correlated to parts of a 4.6 km/sec. twt velocity zone from refraction seismic lines of reid & jackson (1997). they interpreted this zone to include thule supergroup sedimentary rocks, which crop out north of the study area (dawes 1997). thus the upper part of mega-unit h below horizon h1 (fig. 2) probably includes sedimentary rocks with some analogues to the thule supergroup. seismic interpretation suggests that some of the new geophysical and geological mapping of the eastern baffin bay region, offshore west greenland ulrik gregersen, paul c. knutz and john r. hopper 70°w 60°w fig. 2 delta-1 baffin bay 71°n 73°n 75°n 100 km c anada g reenland c anada g reenland shallow well deep exploration well 2d seismic line fig. 1. map of the study area in eastern baffin bay with 2d seismic data used and the location of shallow and deep exploration wells. location of the seismic line of fig. 2 is also shown. 8484 wedge-shaped units in mega-unit h (fig. 2) were probably formed by extension during the proterozoic. sedimentary successions in the baffin bay may also resemble those from the nuussuaq basin, where cretaceous nearshore to deltaic sandy deposits were documented (dam et al. 2009). seismic interpretation below horizon f1 (mega-unit f) shows prograding-aggrading clinoforms and troughs or channels possibly from deltaic systems (gregersen et al. 2013). during parts of the late cretaceous, relative tectonic quiescence prevailed and thick uniform units were deposited in the basins. marine mudstone deposits with suggested source-rock intervals and oil seeps were recovered from both west greenland (bojesen-koefoed et al. 1999) and northern east canadian islands (maclean & williams 1983; brent et al. 2013). within the mapped basins, these marine mudstones may be analogous to deposits in mega-units e and f, between horizons e1 and g1 (fig. 2). extensional faulting occurred during the late cretaceous to earliest paleocene by local rifting in the nuussuaq basin (dam et al. 2009) and locally in a few other places in the region. in north-eastern baffin bay, major cretaceous rift basins trend se–nw and are located east of the kivioq ridge (fig. 4). west of the kivioq ridge, extensive volcanic areas have been mapped primarily from seismic reflection and magnetic anomaly data (fig. 4). studies with refraction seismic data (damm 2010; suckro et al. 2012; altenbernd et al. 2014) showed oceanic crust in the eastern part of the baffin bay basin. seismic facies analyses combined with interpretation of magnetic data in the present study and in gregersen et al. (2013) outline the eastern boundary of the oceanic crust (also named baffin bay volcanics) at nearly the same location (figs 2, 4). the oceanic crust developed as greenland and its rifted margin separated from canada during the paleocene and eocene (oakey & chalmers 2012). the north-east and northward movements of greenland’s rifted basins caused compression-related 2 0 4 t w o -w ay t ra ve l ti m e (s ec .) 6 8 20 km melville bay grabenmbrkirkirbaffin bay basin kib kivioq basin a1 b1 c1 d1 hx h2 a1 quaternary a rchaean c e n o z o i c f2 quaternary sw ne a rchaean c r e t a c e o u s c r e t a c e o u s proterozoic proterozoic c e n o z o i cb1 c1 d1 e1 f1 g1 h1 e1 exex palaeogene baffin bay volcanics (oceanic crust) palaeogene baffin bay volcanics (oceanic crust) f1 f2 f1 e1 f1 e1 g1 h1 hx h2 100 km c anada g reenland 70°w 60°w 73°n 75°n two-way time (ms) 5500 5000 4500 4000 3500 3000 2500 2000 1500 1000 500 fig. 2. seismic line ne–sw across the eastern baffin bay region showing deep basins and large structures. the horizon boundaries of the seismic mega-units a–h are shown by coloured horizons from a1 to hx and likely ages are shown. note that some of the major deep-seated faults continue nearly vertically upwards through parts of the cenozoic section indicating late compressional faulting. the names of the main structural elements are shown below the line and are from sw to ne: baffin bay basin. kir: kivioq ridge. kib: kivioq basin. mbr: melville bay ridge. melville bay graben (fig. 4). the seismic line is courtesy of tgs-nopec geophysical company asa. location of the seismic line is shown in fig. 1. fig. 3. depth-structure map of horizon f1 with two-way travel time. this level is not drilled in the region but interpretation suggests that the horizon may be of a late cretaceous (campanian?) age. the map illustrates the outline of large structures and basins of eastern baffin bay. 85 tectonism with inversion, transtension and thrust faults developing (fig. 2). the palaeogene compression-related faulting in eastern baffin bay was probably a consequence of the same overall plate movements which also caused the eurekan tectonic phase in north greenland and ellesmere island. portions of the large ridge structures likely contain block-faulted and post-rift cretaceous sedimentary rocks later compressed into ridges such as the melville bay ridge and the kivioq ridge (figs 2–4). during the palaeogene, flood basalts and other volcanic rocks developed in the west greenland volcanic province and probably also in the kivioq west volcanic area (fig. 4). paleocene–eocene volcanic rocks cover parts of the cretaceous basins such as in the nuussuaq basin (dam et al. 2009; larsen et al. 2015). samples of flood basalt from the delta-1 well offshore central west greenland (fig. 1) were dated to c. 50–56 ma (nelson et al. 2015). sedimentary geometries suggesting basin-floor fan deposits were formed during the late eocene to oligocene in mega-unit d, above horizon e1 (fig. 2), presumably related to the eurekan compressive phase. this was followed by more passive infilling of the remnant rift basins lasting until middle miocene (horizon d1; fig. 2). during the late miocene to pliocene, large contourite driftand mass-transport systems (e.g. mega-slides) developed in baffin bay within mega-units b and c (between horizons b1 and d1; fig. 2) (knutz et al. 2015). the final phase of basin development is characterised by major progradation of the west greenland shelf towards southwest (mega-unit a between a1 and b1; fig. 2) mainly as a result of trough-mouth fans formed by late pliocene–pleistocene glaciations (knutz et al. 2015). interpretation of seismic facies and attributes has identified a number of amplitude anomalies (bright events) and vertical disturbances (chimney structures) on seismic data over structural highs in the cenozoic section. some of the features could relate to upwards migration of fluids and possible active petroleum systems. some of the source rocks that produce oil seeps in the nuussuaq basin and equivalent sources in west baffin bay could occur in basins of eastern baffin bay (bojesen-koefoed 2011). large structural crests within cretaceous sections occur close to deep basins (fig. 2), and some of the structures may include potential traps for hydrocarbons, given the right conditions. 70°w 60°w 50°w 7 1 °n 7 3 °n 7 5 °n kivioq basin k ivioq ridge w est g reenland volcanic province m elville bay r idge kap york basin k ivioq w est volcanics carey basin upe rn av ik es ca rp men t c anada g reenland m elville bay g raben t t t t t t t t t t t t t t t t t t t t t t t t t t t t upernavik basin n u u ssu aq b asin b a f f i n b a y b a s i n greenland oceanic crust t structural high major sedimentary basin major sedimentary basin and volcanic rock volcanic rock oceanic crust transitional crust 100 km fig. 4. structural elements map of eastern baffin bay. the extent of the canadian part of the baffin bay oceanic crust is from oakey & chalmers (2012). 8686 the updated mapping and new interterpretation in this study have improved the large-scale mapping of the structures and basins of the eastern baffin bay region, outlining a prospectivity potential and also pointing to uncertainties and risks that require future clarification. acknowledgements the study was co-financed by the ministry of mineral resources (government of greenland) and geus. the constructive comments and improvements of the manuscript from referees lynn dafoe and christopher harrison are appreciated and acknowledged. tgs-nopec geophysical company asa is acknowledged for permission to publish the seismic section. references altenbernd, t., jokat, w., heyde, i. & damm, v. 2014: a crustal model for northern melville bay, baffin bay. journal of geophysical research–solid earth 119, 8610–8632. bojesen-koefoed, j.a. 2011: west greenland petroleum systems. an overview of source rocks and oil seepages and their implications for offshore petroleum exploration. danmarks og grønlands geologiske undersøgelse rapport 2011/42, 49 pp. bojesen-koefoed, j.a., christiansen, f.g., nytoft, h.p. & pedersen, a.k. 1999: oil seepage onshore west greenland: evidence of multiple source rocks and oil mixing. in: fleet, a.j. & boldy, s.a.r. (eds): petroleum geology of northwest europe: proceedings of the 5th conference. geological society, london, 305–314. brent, t.a., chen, z., currie, l.d. & osadetz, k. 2013: assessment of the conventional petroleum resource potential of mesozoic and younger structural plays within the proposed national marine conservation area, lancaster sound, nunavut. geological survey of canada, open file 6954, 40 pp. dam, g., pedersen, g.k., sønderholm, m., midtgaard, h., larsen, l.m., nøhr-hansen, h. & pedersen, a.k. 2009: lithostratigraphy of the cretaceous–paleocene nuussuaq group, nuussuaq basin, west greenland. geological survey of denmark and greenland bulletin 19, 171 pp. damm, v. 2010: the expedition of the research vessel “polarstern” to the arctic in 2010 (ark-xxv/3). berichte zur polarund meeresforschung (reports on polar and marine research) 621, 234 pp. alfred wegener institut für polarund meeresforschung, germany, http://hdl.handle.net/10013/epic.36297. dawes, p.r. 1997: the proterozoic thule supergroup, greenland and canada: history, lithostratigraphy and development. geology of greenland survey bulletin 174, 150 pp. gregersen, u., hopper, j.r. & knutz, p.c. 2013: basin seismic stratigraphy and aspects of prospectivity in the ne baffin bay, northwest greenland. marine and petroleum geology 46, 1–18. knutz, p.c., hopper, j.r., gregersen, u., nielsen, t. & japsen, p. 2015: a contourite drift system on the baffin bay–west greenland margin linking pliocene arctic warming to poleward ocean circulation. geology 43, 907–910. larsen, l.m., pedersen, a.k., tegner, c., duncan, r.a., hald, n. & larsen, j.g. 2015: age of tertiary volcanic rocks on the west greenland continental margin: volcanic evolution and event correlation to other parts of the north atlantic igneous province. geological magazine 153(3), 487511, http://dx.doi.org/10.1017/s0016756815000515. maclean, b. & williams, g.l. 1983: geological investigations of baffin island shelf in 1982. in: current research, part b, geological survey of canada, paper 83-1b, 309–315. nelson, c.e., jerram, d.a., clayburn, j.a.p., halton, a.m. & roberge, j. 2015: eocene volcanism in offshore southern baffin bay. marine and petroleum geology 67, 678–691. oakey, g.n. & chalmers, j.a. 2012: a new model for the paleogene motion of greenland relative to north america: plate reconstructions of the davis strait and nares strait regions between canada and greenland. journal of geophysical research 117, 1–28. reid, i. & jackson, h.r. 1997: crustal structure of northern baffin bay: seismic refraction results and tectonic implications. journal of geophysical research 102, 523–542. rolle, f. 1985: late cretaceous – tertiary sediments offshore central west greenland: lithostratigraphy, sedimentary evolution, and petroleum potential. canadian journal of earth sciences 22, 1001–1019. suckro, s.k., gohl, k., funck, t., heyde, i., ehrhardt, a., schreckenberger, b., gerlings, j., damm, v. & jokat, w. 2012: the crustal structure of southern baffin bay: implications from a seismic refraction experiment. geophysical journal international 190, 37–58. whittaker, r.c., hamann, n.e. & pulvertaft, t.c.r. 1997: a new frontier province offshore northwest greenland: structure, basin development, and petroleum potential of the melville bay area. american association of petroleum geologists bulletin 81, 978–998. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark; e-mail: ug@geus.dk geological survey of denmark and greenland bulletin 35, 2016, 27-30 27© 2016 geus. geological survey of denmark and greenland bulletin 35, 27–30. open access: www.geus.dk/publications/bull as part of its strategy, the geological survey of denmark and greenland (geus) is to develop a national, digital 3d geological model of denmark that can act as a publicly accessible database representing the current, overall interpretation of the subsurface geology. a national model should be under constant development, focusing on meeting the current demands from society. the constant improvements in computer capacity and software capabilities have led to a growing demand for advanced geological models and 3d maps that meet the current technical standards (berg et al. 2011). as a consequence, the users expect solutions to still more complicated and sophisticated problems related to the subsurface. geus has a long tradition of making 2d maps of subsurface layer boundaries and near-surface geology (fredericia & gravesen 2014), but in the change from 2d to 3d and when combining data in new ways, new geological knowledge is gained and new challenges of both technical and organisational character will arise. the purpose of this paper is to present the strategy for the national 3d geological model of denmark and the planned activities for the years ahead. the paper will also reflect on some of the challenges related to making and maintaining a nationwide 3d model. initially, the model will only include the danish onshore areas, with the danish offshore areas and greenland to be added later using a similar general setup. the elements of the national 3d geological model the national model will be constructed as a 3d geological framework model consisting of a number of surfaces. these surfaces can represent top and bottom of defined geological formations, stratigraphic complexes or other types of spatially recognisable units, as well as erosional surfaces, stratigraphic markers or transgressional surfaces. the surfaces will characterise units defined in a legend of the danish subsurface compiled in connection with the national 3d geological model project. interpretation points, lines, polygons, etc. will define the surfaces together with an interpolated or triangulated grid. this framework model is planned to grow continually with the addition of new surfaces. the layered framework model will be supplemented by volumetric cells containing detailed geological information between mapped surfaces (fig. 1). lithology and lithofacies or related parameters, such as porosity or resistivity, can be added as attributes to the cells. the model is intended to contain varying levels of detail and it will be possible to use the model at different scales. the users will be able to download the exact elements they require for a specific modelling purpose within a particutowards a national 3d geological model of denmark peter b.e. sandersen, thomas vangkilde-pedersen, flemming jørgensen, richard thomsen, jørgen tulstrup and johnny fredericia fig. 1. model example combining layers and volumetric cells. the manmade layers in the uppermost part of the subsurface are represented as volumetric pixels (voxels), and the deeper parts of the succession as layers. based on the 3d model of the municipality of odense (mielby et al. 2015). 2828 lar area, i.e. near-surface layers for further hydrogeological modelling in connection with groundwater resource management. this will, of course, require guidance regarding the scale in which the extracted model element can be used. standards and procedures will be developed for the construction of the model elements as well as for performing quality assurance, quality assessment and model updates. the national model is planned to be platformindependent and by using standardised import and export formats, the construction of model elements will therefore not be restricted to specific modelling software packages. the model will initially contain a number of key surfaces, but details will continuously be added. at present, a number of surfaces encompassing the deep succession from top pre-zechstein to the top of the chalk group exist in preliminary versions (table 1). above this part of the succession, construction of 1–2 major tertiary surfaces and the top of the pre-quaternary surface is planned. modelling of the complex quaternary geology will be done locally and regionally by using layer boundaries and volumetric cells. because the amount of dated quaternary sediments in denmark is limited these model elements will not necessarily be modelled as specific lithoor chronostratigraphic units, but as units solely defined by lithology. the model database constitutes an important part of the national 3d geological model containing all the required model elements while at the same time being sufficiently flexible when future elements are added. use of existing data and geological models geus hosts a wide range of publicly accessible databases (e.g. ditlefsen et al. 2012; hansen & pjetursson 2011; møller et al. 2009a, b; tulstrup 2004). the most significant databases are the jupiter well database, the geophysical database gerda, the oil and gas database frisbee and the geological models database. these databases will constitute the main data supply for the national 3d geological model, and therefore their continuous update is very important. the incorporation into the national 3d geological model will ensure that the vast amount of data collected over more than a century is continually used and updated. in denmark, 2d and 3d geological mapping and modelling have been performed regionally and locally primarily by geus, universities, regional authorities and municipalities, oil companies, the mining industry and consulting companies. models have typically been made in connection with oil and gas exploration, groundwater and drinking water projects, geothermal projects, raw material and mineral exploration as well as soil and groundwater contamination issues. especially the intensive groundwater mapping campaign over the past 15 years has produced a large number of publicly available models for the upper parts of the subsurface (thomsen et al. 2013). this has resulted in a patchwork of 2d and 3d models, but because the models have been built in different ways they are difficult to merge. therefore, in the process of making a new national 3d geological model it is necessary to evaluate and amend the existing models. the national 3d geological model will be dynamic and regularly updated in order to include new data and interpretations. the modelling will therefore follow strict versioning procedures, with options for displaying the data on which updated interpretations are based. model uncertainties, quality assurance and quality control handling of model uncertainty has been discussed intensely for many years and many different suggestions have been put forward (e.g. lark et al. 2014; sandersen 2008; tacher et al. 2006; wellmann et al. 2014). the change from 2d to 3d and the ensuing increased model complexity will create a demand from users for assessment of the inherent uncertainty of the model. this issue is highly complex and the challenges are numerous. the uncertainty of a geological model will be a mixture of the uncertainties of each dataset and the uncertainties of the interpretations of the combined dataset, with the additional challenge of handling both the quantitative and the qualitative aspects. model element surface period pre-quaternary succession top table 1. planned main surfaces in the national 3d geological model neogene and older palaeogene succession top palaeogene top bottom top bottom top bottom top bottom top bottom top bottom chalk group frederikshavn formation haldager sand formation fjerritslev formation gassum formation bunter sandstone formation zechstein group top bottom triassic zechstein cretaceous cretaceous/jurassic jurassic jurassic triassic 29 a concept for assessing the uncertainties of the individual model elements will be made specifically for the national 3d geological model to ensure that all model elements are evaluated in the same way. the uncertainty concept will not necessarily be adaptable for use in models other than the national 3d geological model, because it will be tailored specifically to the chosen model setup. the national 3d geological model will be disseminated to a variety of users and therefore a transparent construction process and a thorough documentation are required. descriptions of workflows, procedures and guidelines for quality assurance and quality control will be handled by an editorial function in order to keep up a high degree of consistency. discussion in general, varying data coverage and data resolution pose a major challenge for the construction of the national 3d geological model. in the deep parts of the subsurface (below 300–400 m) the data are primarily from boreholes and seismic surveys carried out by the oil and gas industry and with a coverage dictated by the focus of the exploration surveys. in the shallow parts of the subsurface, the data are primarily collected in connection with groundwater investigations, raw material or mineral exploration, soil contamination investigations and geotechnical projects. data originate from boreholes, seismic surveys, electric and electromagnetic surveys and outcrops. generally the data density is much higher than in the deeper parts of the subsurface. the shallow part of the danish subsurface is greatly affected by pleistocene glaciations. this part is therefore particularly lithologically and structurally complex and requires a large amount of detailed input data. even though the data density is rather high in the uppermost parts of the subsurface, the data are in fact often geographically clustered. therefore, in many areas detailed modelling can only be made within such data clusters. in reality, the result is a patchwork of areas with a high data density in the survey focus areas and a low density in the surrounding areas. this means that the varying data density will be reflected in the geological model interpretations resulting in certain geological elements not being resolved outside the data clusters (fig. 2). this challenge can be met by constructing a model that can handle different scales with different degrees of detail, but not necessarily with a full geographical coverage of the most detailed interpretations. a model like this will display the actual status of the mapping and show where the geological knowledge is good and where not. in modern 3d modelling software there are no zoom limitations, and if not otherwise stated, the data and the interpreted model can be viewed and evaluated at any scale. it is therefore important to convey the scalerelated limitations of the model to the users. the potential users of the national 3d geological model are numerous (i.e. waterworks, municipalities, governmental agencies, raw-materials and minerals industry, private consultants and educational institutions), and they will have a variety of purposes for their use of the downloaded model elements. to cater for as many of these as possible, the model will be constructed as a multi-purpose tool. this means that the output from the model will require standardised, off-the-shelf products as well as individually tailored elements based on the same framework. an example of a standardised product from the ‘3d model department store’ could be a suite of nation-wide, fixed-scale surfaces to be used in projects dealing with regional or national assessments. a tailored product from the ‘3d custom shop’, on the other hand, could be the delivery of a number of specific surfaces in a selected area supplied with lithological information in volumetric cells to be used in a geographically small-scale project. in this way the 3d model construction procedures can be kept stringent, while the model output can be more flexible in order to meet specific user needs. high data density (boreholes, tem data) low data density (boreholes) 1 km ? ? ? 100 m fault 0 –100 fig. 2. cross-section sketch illustrating the challenges of modelling in areas with low data density. vertical rods represent boreholes with clay (orange) and sand (red). coloured background in the area with high data density shows measured and gridded sediment resistivity values (tem data; transient electromagnetic method). red and purple: high resistivities. g reen and blue: low resistivities. 3030 summary and perspectives the national 3d geological model of denmark will be constructed from the geological and geophysical data collected during decades of surveying. it will also be based on existing 2d maps and geological models, from which specific elements will be extracted and incorporated. all basic data used in the national 3d geological model are stored in an array of geus databases. the model will be dynamic and continuously updated in order to maintain relevance and appeal to the end users. the model will provide a wide range of end users with standardised model downloads as well as with tailor-made products. building a national 3d geological model for the country will require a considerable effort over a long period and therefore the model is planned to start with only a limited number of key elements. it will then continue to grow and in the process seek to adapt to the varying demands from society. the short-term strategy is to establish at least 15 key surfaces within the next four years and initiate construction of regional and local geological elements in the shallow parts of the model. the model databases and a beta-version of the web interface will be established and launched. in addition, standards and procedures related to the construction of the 3d model will be described in a series of guidelines. the long-term strategy is to include local and regional surfaces of the main surveyed areas in the national model within the next 10 years, including the danish offshore areas. an established dialogue with potential users throughout the process will help target the model contents towards a versatile national model that meets the requirements of the public and at the same time secures the use of the huge amount of valuable data collected over several decades. acknowledgements the authors would like to thank the large number of geologists who have contributed to and participated in the workshops and discussion groups during the initial phases of the work on the national 3d geological model. references berg, r.c., mathers, s.j., keefer, d.a. & kessler, h. (eds) 2011: a synopsis of current three-dimensional geological mapping and modeling in geological survey organizations. illinois state geological survey open file report, circular 578, 2011. ditlefsen, c., gausby, m., salomonsen, j. & hansen, m. 2012: vejledning i anvendelse af modeldatabasen. geovejledning 9, geus special publication. 31 pp. fredericia, j. & gravesen, p. 2014: 125 years of geological research for society. geological survey of denmark and greenland bulletin 31, 9–14. hansen, m. & pjetursson, b. 2011: free, online danish shallow geological data. geological survey of denmark and greenland bulletin 23, 53–56. lark, r.m., thorpe, s., kessler, h. & mathers, s.j. 2014: interpretative modelling of a geological cross section from boreholes: sources of uncertainty and their quantification. solid earth 5, 1189–1203. mielby, s. et al. 2015: udvikling af en 3d geologisk/hydrogeologisk model som basis for det urbane vandkredsløb. syntese rapport. geological survey of denmark and greenland special publication (in danish only). møller, i., søndergaard, v. h., jørgensen, f. auken, e. & christiansen, a.v. 2009a: integrated management and utilization of hydrogeophysical data on a national scale. near surface geophysics 7, 647–659. møller, i., søndergaard, v.h. & jørgensen, f. 2009b: geophysical methods and data administration in danish groundwater mapping. geological survey of denmark and greenland bulletin 17, 41–44. sandersen, p.b.e. 2008: uncertainty assessment of geological models – a qualitative approach. in: refsgaard, j.c. et al. (eds): calibration and reliability in groundwater modelling: credibility of modelling. international association of hydrological sciences publication 320, 345–349. tacher, l., pomian-srzednicki, i. & parriaux, a. 2006: geological uncertainties associated with 3-d subsurface models. computers & geosciences 32, 212–221. thomsen, r., sondergaard, v. & klee, p. 2013: greater water security with groundwater – groundwater mapping and sustainable groundwater management. the rethink water network and danish water forum white papers, copenhagen. available at www.rethinkwater.dk. tulstrup, j. 2004: environmental data and the internet: openness and digital data management. geological survey of denmark and greenland bulletin 4, 45–48. wellmann, f.j., lindsay, m., poh, j. & jessell, m. 2014: validating 3-d structural models with geological knowledge for improved uncertainty evaluations. energy procedia 59, 374–381. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark; e-mail: psa@geus.dk geological survey of denmark and greenland bulletin 13, 2007, 33-36 the western margin of the greenland craton has been much less stable in the phanerozoic than previously thought. this new insight has come from close integration of independent data sets: geomorphological analysis of large-scale landscapes, apatite fission track analysis (afta), onshore and offshore stratigraphy and analysis of onshore fault and fracture sys tems. each data set records specific and unique parts of the event chronology and is equally important to establish a consistent model. a key area for understanding the mesozoic– cenozoic landscape evolution and into the present is the uplifted part of the nuussuaq basin, where remnants of planation surfaces cut across the cretaceous to eocene sedimentary and volcanic rocks. our integrated analysis concluded that the west greenland mountains were formed by late neogene tectonic uplift (fig. 1) and also provided new insight into early phanerozoic development. to understand our model, we present the different methods and the results that can be deduced from them. basic concepts the mapping of volcanic and sedimentary successions within the nuussuaq basin is crucial for understanding the late mesozoic–palaeogene landscape development (e.g. dam et al. 1998; chalmers et al. 1999; dalhoff et al. 2003). especially important for the landscape analysis is the availability of maps showing vertical geological sections (pedersen et al. 2006). exploration for hydrocarbons has resulted in many seismic data, and several deep wells have been drilled both onshore and offshore (e.g. chalmers et al. 1999; piasecki 2003). landscape analysis aims at setting up a relative tectonic event chronology through identification and mapping of both extensive baselevel governed surfaces and re-exposed surfaces. these palaeosurfaces, formed by erosion in climates or tectonic settings different from the present, cut across bedrock of different ages and can be arranged in chronological order based on (1) stratigraphical relationships with cover rocks, (2) geometrical relationships between different palaeosurfaces, and (3) analysis of the detailed forms of the largescale landscapes, reflecting climatic-driven formation processes. the baselevel is fundamental, as lowering of baselevel (an uplift event) causes valley incision and initiation of surface formation while raising baselevel (subsidence) causes palaeosurfaces to be preserved below cover rocks (bonow 2005; bonow et al. 2006a, b). afta is a method for defining the temperature history of rock samples, based on analysis of radiation damage features (‘fission tracks’) produced by spontaneous fission of 238u atoms within apatite crystals. tracks are produced continu© geus, 2007. geological survey of denmark and greenland bulletin 13, 33–36. available at: www.geus.dk/publications/bull 33 a multi-disciplinary study of phanerozoic landscape development in west greenland johan m. bonow, peter japsen, paul f. green, robert w. wilson, james a. chalmers, knud erik s. klint, jeroen a.m. van gool, karna lidmar-bergström and asger ken pedersen fig. 1. left: study area with precambrian basement and cover rocks that are crucial for determining the relative age of palaeosurfaces. the rose diagram summarises the regional lineament patterns, based on field mapping in the framed area (cf. wilson et al. 2006). the relationship between lineaments onshore and offshore allows for a relative event chronology; colouring refers to timing (cf. fig. 3). position of gro#3 well indicated. es, etch surface of late mesozoic – paleocene age. note its position close to cover rocks. modified from bonow et al. (2006a). right: topography. a regionally developed oligo cene–miocene planation surface was differentially uplifted on separate tectonic blocks in the neogene. today it is close to the summit level and has been tilted in different directions. ously over geological time, but are shortened at a rate that depends on the prevailing temperature, until at temperatures higher than c. 120°c tracks are totally erased (‘annealing’). fission track age and track length data provide the basis for estimating the time at which a sample began to cool from a palaeo-thermal maximum as well as the magnitude of the maximum palaeotemperature. cooling can be interpreted as either change of heat-flow within the crust or erosion of overlying rocks (e.g. green et al. 2002). structural analysis of faults and fracture systems aims to establish a relative chronology of tectonic movements that have changed the stress field, as a change will lead to the formation of a new set of faults and possibly the reactivation of older ones. structural analysis of the area between nuussuaq and sisimiut forms the basis for a regional model, explaining different tectonic movements through time (wilson et al. 2006). key results geology. the cretaceous–palaeogene sedimentary and volcanic successions within the nuussuaq basin record deep incision of valleys in the maastrichtian and early paleocene (e.g. dam et al. 1998), subsidence during volcanism, and deposition of marine sediments within the volcanic succession now at high elevation (piasecki et al. 1992), which are evidence that both uplift and subsidence of kilometre scale took place during and after rifting (chalmers et al. 1999). during the palaeogene the basalts offshore (and probably onshore, japsen et al. 2006a) became buried below sediments. seismic sections west of nuussuaq show that palaeogene and younger sequences have been tilted seawards and truncated at a late date (chalmers 2000). 34 fig. 2. views of the etch surface (es) and the upper and lower planation surfaces (ups and lps) formed in basement rocks. a: the re-exposed es at fortunebay, southern disko. the white line shows the approximate border between paleocene basalt and gneiss. area location in fig. 1. photo location in b. b: 3d model of the fortunebay area. note the oligocene–miocene planation surface at high elevation across the basalt. modified from bonow (2005). c: the stripped es at nassuttooq. location in fig 1. d: the oligocene–miocene ups east of sukkertoppen iskappe cuts across precambrian basement. location in e. e: 3d model showing the well-preserved ups east of sukkertoppen iskappe. red frame indicates position of map in f. f: map showing ups and lps. hill complexes rising above the ups may be part of a sub-ordovician peneplain. modified from bonow et al. (2006a). geomorphology. three different palaeosurfaces in the pre cam brian basement have been identified in west greenland, viz. a surface formed by deep weathering and stripping of the weathering mantle (etch surface, es), and an upper and lower planation surface (ups and lps; bonow 2005; bonow et al. 2006a, b). the es is characterised by distinct hills (fig. 2a) and re ceived its final shape in part prior to the deposition of upper cretaceous deltaic sediments and in part prior to the extrusion of palaeogene basalts (fig. 2b). the es can mainly be identified at low elevations and close to cover rocks (figs 1, 2c). the ups has low relative relief compared to the es (fig. 2d) and must be younger as it cuts across both mideocene basalts and the etch surface. the ups forms the summits of differentially tilted, fault-bounded tectonic blocks. a planation surface cannot be formed as an inclined plain because any tilt would cause valleys to incise and the relief to rejuvenate towards the baselevel (bonow et al. 2006b, fig. 6). the lps was formed in response to lowered baselevel (uplift) and became incised into the ups (fig. 2e). furthermore, summits of distinct hill complexes above the ups (fig. 2f) may relate to a sub-ordovician palaeosurface because remnants of lower palaeozoic rocks suggest that west greenland may have had a long-lasting palaeozoic cover (bonow et al. 2006a). consequently, erosion of precambrian basement rocks has been limited since the early palaeozoic, but this does not ex clude deposition and subsequent removal of thick sequences of phanerozoic cover rock as indicated by afta data. thermochronology. afta data from cretaceous sedimentary rocks define three major cenozoic cooling episodes, while basement samples define major triassic and jurassic cooling episodes (related to rifting?), and also earlier (palaeo zoic) episodes (table 1). the deepest samples in the 3 km deep gro#3 well on nuussuaq are totally annealed (fig. 1; japsen et al. 2005). the progressive development of fission tracks can therefore be followed through the sedimentary section, giving a rare opportunity to resolve the details of the late ceno zoic cooling history. oligocene cooling involved both exhu mation and a decrease in basal heat flow, while miocene and pliocene cooling episodes were dominantly related to exhu mation. the two latest cooling events constrain the cooling events into the present. faults and fractures. analysis of regional lineament trends shows five main systems that fit a two-stage model (wilson et al. 2006). a system of n–sand nnw–sse-trending normal faults reflects the fault patterns in the davis strait during the late cretaceous to paleocene. this system was overprinted and reactivated by strike-slip faults associated with a later nne–ssw-trending sinistral wrench system that re flects the development of the ungava transform system during the eocene. 35 fig. 3. event chronology for central west greenland based on data from the separate disciplines. integration of these data sets shows that the western margin of the greenland craton has been less stable than previously thought. shaded intervals indicate the proven age of geological units, onshore and offshore, and the estimated time required for formation of palaeosurfaces. horizontal arrows indicate extensional reactivation and opening of vertical fault and fracture systems. the associated intervals indicated by stippled lines illustrate the probable geological age range and uncertainties in dating palaeo-surfaces. es, etch surfaces; ups, upper planation surface; lps, lower planation surface. data from piasecki et al. (1992); dam et al. (1998); chalmers et al. (1999); chalmers (2000); dalhoff et al. (2003); piasecki (2003); bonow (2005); japsen et al. (2005, 2006a); bonow et al. (2006a, b, c); pedersen et al. (2006) and wilson et al. (2006). the model and future implementation our model shows where in time independent constrained data exist and time-frames for uncertainties and lack of data (fig. 3). the model shows that each discipline has long periods of no data, but when combined only few periods have no data representation at all. in particular, landscape analysis and afta data complement each other, because palaeosurfaces show that rock was exposed at the landsurface, whereas afta data indicate when and by how much a palaeosurface has been buried. this approach shows that the present summits were buried below up to 1 km of rocks prior to eocene–oligocene uplift, and that the ups formed during the oligocene– miocene due to stable baselevel conditions. similarly, uplift in the late miocene resulted in valley incision (the lps) and tilting of the ups. final uplift in the ?pliocene resulted in the present-day mountains. late uplift reactivated and opened the fault and fracture systems, thus facilitating both weathering and the development of a coastal escarpment (bonow et al. 2006c). our model is also used in ongoing uplift studies in south-west greenland (japsen et al. 2006b). the integration of data from geomorphology, thermo chrono logy, geology and fault/fracture patterns to show that the present landscape of west greenland is the result of tectonic movements throughout the phanerozoic with significant movements also in the neogene and even into the present (fig. 3). the approach presented here may be applied to understand landscape development along other passive continental margins. acknowledgements this work was supported by the carlsberg foundation, the bureau of minerals and petroleum, the danish natural science research council, the swedish research council, arktisk station, stiftelsen margit althins stipendiefond, svenska sällskapet för antropologi och geografi and john söderbergs stiftelse. references bonow, j.m. 2005: re-exposed basement landforms in the disko region, west greenland – disregarded data for estimation of glacial erosion and uplift modelling. geomorphology 72, 106–127. bonow, j.m., lidmar-bergström, k. & japsen, p. 2006a: palaeosurfaces in central west greenland as reference for identification of tectonic movements and estimation of erosion. global and planetary change 50, 161–183. bonow, j.m., japsen, p., lidmar-bergström, k., chalmers, j.a. & pedersen, a.k. 2006b: cenozoic uplift of nuussuaq and disko, west greenland – elevated erosion surfaces as uplift markers of a passive margin. geomorphology 80, 325–337. bonow, j.m., klint, k.e.s. & japsen, p. 2006c: the nordre isortoq escarpment. field report summer 2005. danmarks og grønlands geologiske undersøgelse rapport 2006/13, 68 pp. chalmers, j.a. 2000: offshore evidence for neogene uplift in central west greenland. global and planetary change 24, 311–318. chalmers, j.a., pulvertaft, c., marcussen, c. & pedersen, a.k. 1999: new insight into the structure of the nuussuaq basin, central west green land. marine and petroleum geology 16, 197–224. dalhoff, f., chalmers, j.a., gregersen, u., nøhr-hansen, h., rasmussen, j.a. & sheldon, e. 2003: mapping and facies analysis of paleocene – mid-eocene seismic sequences, offshore southern west greenland. marine and petroleum geology 20, 935–986. dam, g., larsen, m. & sønderholm, m. 1998: sedimentary response to mantle plumes: implications from paleocene onshore successions, west and east greenland. geology 26, 207–210. green, p.f., duddy, i.r. & hegarty, k.a. 2002: quantifying exhumation from apatite fission-track analysis and vitrinite reflectance data: precision, accuracy and latest results from the atlantic margin of nw europe. in: doré, a.g. et al. (eds): exhumation of the north atlantic margin: timing, mechanisms and implications for petroleum exploration. geological society special publication (london) 196, 331–354. japsen, p., green, p.f. & chalmers, j.a. 2005: separation of palaeogene and neogene uplift on nuussuaq, west greenland. journal of the geological society (london) 162, 299–314. japsen, p., bonow, j.m., green, p.f., chalmers, j.a. & lidmar-bergström, k. 2006a: elevated, passive continental margins: long-term highs or neogene uplifts? new evidence from west greenland. earth and planetary science letters 248, 315–324. japsen, p., bonow, j.m., peulvast, j.-p. & wilson, r.w. 2006b: uplift, erosion and fault reactivation in southern west and south greenland. field report summer 2006. danmarks og grønlands geologiske undersøgelse rapport 2006/63, 77 pp. pedersen, a.k., larsen, l.m., pedersen, g.k. & dueholm, k.s. 2006: five slices through the nuussuaq basin, west greenland. geological survey of denmark and greenland bulletin 10, 53–56. piasecki, s. 2003: neogene dinoflagellate cysts from davis strait, offshore west greenland. marine and petroleum geology 20, 1075–1088. piasecki, s., larsen, l.m., pedersen, a.k. & pedersen, g.k. 1992: palynostratigraphy of the lower tertiary volcanics and marine clastic sediments in the southern part of the west greenland basin: implications for the timing and duration of the volcanism. rapport grønlands geologiske undersøgelse 154, 13–31. wilson, r.w., klint, k.e.s., van gool, j.a.m., mccaffrey, k.j.w., holdsworth, r.e. & chalmers, j.a. 2006: faults and fractures in central west greenland: on-shore expression of continental break-up and sea-floor spreading in the labrador – baffin bay sea. geological survey of denmark and greenland bulletin 11, 185–204. authors’ addresses j.m.b., p.j., j.a.c., k.e.s.k. & j.a.m.v.g., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: jbon@geus.dk p.f.g., geotrack international, 37 melville road, brunswick west, victoria 3055, australia. r.w.w., reactivation research group, department of earth sciences, university of durham, durham, dh1 3le, uk. k.l.-b., department of physical geography and quaternary geology, stockholm university, se-106 91 stockholm, sweden. a.k.p., geological museum, university of copenhagen, øster voldgade 5–7, dk-1350 copenhagen k, denmark. 36 geological survey of denmark and greenland bulletin 35, 2016, 91-94 91© 2016 geus. geological survey of denmark and greenland bulletin 35, 91–94. open access: www.geus.dk/publications/bull the continental shelf of labrador and newfoundland has a long history of hydrocarbon exploration, and the accumulated oil production from the northern grand banks exceeds one billion barrels (fig.1). the canada-newfoundland & labrador offshore petroleum board (www.cnlopb.ca) awarded several new licenses on the northern grand banks in 2015 and announced licensing rounds for the labrador sea region in the coming years. vertical motion along passive continental margins such as the atlantic margin of canada, plays an essential role in shaping these margins and their petroleum systems; in particular by removing sedimentary strata (japsen et al. 2012; green et al. 2013). it is thus a fundamental question whether a hiatus in the stratigraphic record represents an episode of stability and non-deposition or an event involving deposition followed by removal of rocks. in this context, a hiatus represents not only a gap in the stratigraphic record, but also a gap in our understanding of the geological history. in terms of hydrocarbon systems, failure to account for greater depths of burial prior to exhumation can lead to serious underestimation of the maturity of petroleum resources. similarly, the effects of exhumation on the timing of hydrocarbon generation, on changes in migration routes and on any reservoir hydrocarbons also require assessment (doré et al. 2002). insights into the uplift history of a margin are also important for understanding the source-tosink system of sediment input into offshore basins. in broader terms of geological development, understanding the history of vertical movements along a passive continental margin is important for investigating whether the elevated regions along these margins, such as the torngat mountains in northern labrador (fig. 1), are either (a) the eroded remnants of ancient orogens (mcgregor et al. 2013), (b) rift shoulders related to processes during rifting and break-up (weissel & karner 1989) or (c) the results of postbreakup episodes of burial and exhumation driven by platetectonic forces (japsen et al. 2006, 2012; green et al. 2013). studies of the burial and exhumation history both onshore and offshore of labrador and newfoundland are, however, scarce, but several observations – that we review in the following – indicate that a number of uplift episodes followed by denudation, both preand post-breakup, shaped the present-day structure of the margin. evidence for episodes of burial and exhumation of the margin the atlantic margin of canada has many features in common with passive continental margins in other parts of the world (japsen et al. 2012; green et al. 2013), such as elevated plateaux (i.e. regional high-level landscapes of low, relative relief) at 1 to 2 km or more above sea level (a.s.l.) cut by deeply incised valleys and commonly separated from an adjacent coastal plain by one or more escarpments. the torngat mountains with peaks reaching 1.7 km a.s.l. in northern labrador, slope much more steeply towards the labrador sea than they do towards the hinterland farther burial and exhumation history of the labradornewfoundland margin: first observations peter japsen, paul f. green, johan m. bonow, alana m. hinchey and derek h.c. wilton afta outcrop sample afta sample, pilot study afta drill hole sample elevation (km) 0 1 2-1-2-3-4-5 250 km labrador newfoundlandnewfoundland st. john’s grand banks 60°w65°w 50°n 60°n t o rngat m o untains cretaceous outlier g reenland c an ad a fig. 2b fig. 3 hopedale e-33hopedale e-33 fig. 2a nova scotia fig. 1. outline of the study area, c. 1500 km along the labrador-newfoundland margin between 46 and 60.5°n. yellow star: cretaceous outlier at schefferville (dorf 1967). afta: apatite fission-track analysis. 9292 west in interior labrador, and their overall shape is thus similar to that of the coastal mountains in greenland (fig. 2). however, farther south (e.g. on newfoundland) elevations do not reach 1 km a.s.l. as along other passive margins, mesozoic–cenozoic rift systems parallel the labrador-newfoundland margin with a transition from continental to oceanic crust farther offshore. here breakup occurred in the early cretaceous east of newfoundland and in the paleocene east of labrador. the synand post-rift sediments at the landward margin of these rifts dip towards the rifts and are truncated by one or more shallow unconformities or by the seabed (fig. 3). in particular, the margin of labrador shares the characteristics listed above with the conjugate margin of west greenland where the geological record documents that the present-day high mountains are not remnants of the rifting process but the result of much later uplift which partially removed thick, post-rift deposits (japsen et al. 2006). evidence from the offshore domain dickie et al. (2011) noted that the tertiary sediments along labrador are tilted seaward and truncated (fig. 3), and that late oligocene as well as younger (possibly miocene) unconformities might correspond to phases of uplift of the labrador margin as proposed by mcmillan (1973). according to dickie et al. (2011), the sedimentary record along the labrador margin is difficult to interpret because of the limited dating of the younger, post-oligocene section, and the many phases of channelling and erosion that are exhibited. subsequently, ainsworth et al. (2014) studied the cretaceous–tertiary stratigraphy of the labrador shelf and provided improved constraints on several unconformities; in particular they documented the presence of a regional miocene hiatus. figure 4 shows evidence that the pre-pliocene sequences along labrador have been more deeply buried in the past, most likely prior to the removal of miocene strata. also on the grand banks, uplift and exhumation were important processes for the shaping of petroleum systems (sinclair et al. 1994; avery 2001). evidence from the onshore domain precambrian rocks dominate labrador and newfoundland, but phanerozoic cover rocks are present across the region, for example in the palaeozoic basins of western newfoundland (cooper et al. 2001). hendriks et al. (1993) interpreted apatite fission-track data from western newfoundland to indicate major episodes of late carboniferous and jurassic exhumation, which agree with evidence from thermal maturity of palaeozoic rocks that they had been more deeply buried below a cover up to 3 km thick (wila b fig. 2. similar landscapes on the conjugate margins across the labrador sea. a: elevated plain (c. 800 m a.s.l.) across precambrian basement cut by a deep valley, torngat mountains, labrador. b: elevated plain (c. 900 m a.s.l.) across paleocene basalts cut by a deep valley, disko island, west greenland. the study will investigate whether the elevated plain is a cenozoic erosion surface as it is the case for the conjugate margin in west greenland (bonow et al. 2006; japsen et al. 2006). photo locations in fig. 1. 5 km sw ne 1 1 2 3 4 5 2’ 2 t w t ( s) 3 hopedale e-33 late eocene marker fig. 3. seismic profile off labrador illustrating post-eocene tilting and truncation of the sedimentary sequences (after dickie et al. 2011). unconformities: 5: late eocene. 4: base eocene. 3: mid-paleocene. 2’: late cretaceous. 2: mid-cretaceous. 1: top basement. sand-prone units: gold and yellow colours. location on fig. 1. twt: two-way travel time. 93 liams et al. 1998). pe-piper & mackay (2006) presented evidence for early cretaceous drainage from western newfoundland to areas onshore and offshore nova scotia. an outlier of cretaceous sediments on central labrador around schefferville (dorf 1967) constrains the construction of a relative denudation chronology based on stratigraphic landscape analysis and for thermochronological modelling by documenting when basement rocks were exhumed to the surface. white et al. (2000) included the schefferville outlier in their evidence for an albian connection between the labrador sea and the cretaceous western interior seaway of north america. grist & zentilli (2003) reported evidence for post-jurassic exhumation with as much as 30°c of post-paleocene cooling of the southern portion of the canadian atlantic margin, based on fission-track data from e.g. nova scotia. they also reported high vitrinite reflectance values for jurassic strata in the fundy basin, offshore nova scotia, in agreement with 2 km of post-jurassic erosion inferred from seismic data. integrated investigation of the vertical movements along the margin the observations reviewed above demonstrate that episodes of burial and exhumation have affected the atlantic margin of canada both prior to and after break-up. the available evidence does not, however, allow further definition of the timing and magnitude of the vertical movements that shaped the present-day margin. in particular, it is not possible to define when the mountains along the margin reached their present elevation. we have therefore initiated a research project aimed at defining the main events of burial and exhumation along the margin onshore and offshore labrador and newfoundland (fig. 1). the study has three components: (1) a thermochronological study based on samples from outcrops and from onshore and offshore boreholes with associated thermal history interpretations (green et al. 2013). a pilot study comprising apatite fission-track analysis (afta) data in 12 samples (fig. 1) revealed a long history of phanerozoic cooling and exhumation episodes, notably a regional triassic event during which a sample of latest neoproterozoic sandstone collected in st. john’s cooled below 110°c, corresponding to the onset of removal of a kilometre-thick cover of palaeozoic–triassic rocks. (2) a stratigraphic landform analysis of the study area based on digital elevation data and stratigraphic information to map exposed denudation surfaces. this analysis provides evidence of both uplift and subsidence using cross-cutting relationships between palaeosurfaces (onshore unconformities expressed as large-scale, low-relief surfaces produced by erosion to base level) and stratigraphic constraints. we intend to use this analysis to construct a relative chronology for surface formation and tectonic events (green et al. 2013). (3) an integrated interpretation of the geological, geomorphological and thermochronological data. we will combine the relative denudation chronology from the stratigraphic landscape analysis with the absolute timing of cooling events determined from the afta data in order to estimate the timing and magnitude of uplift and exhumation along the margins of labrador and newfoundland. studying uplift/exhumation with just one technique in isolation provides only part of the story. without the afta data, the landform analysis will only yield a relative event chronology. without the landform analysis, it will not be possible to conclude whether surface uplift accompanied exhumation as recorded by the afta data. summary the stratigraphic record along the continental margin of labrador and newfoundland provides ample evidence for quaternary saglek mokami kenamu cartwright markland precambrian vr values from robertson research vr values from avery 2008 vr values from bujak davies group vr values from geochem maturity (%ro) 1.00.50.3 0.70.2 0.4 750 m deeper burial 1.5 1 0.5 0 d ep th ( km b el o w s ea b ed ) hopedale e-33 possible reconstruction dth water depth 550 m fig. 4. scattered vitrinite reflectance (vr) data from various labs for the hopedale e-33 well (source: canada basin database; http://basin. gdr.nrcan.gc.ca/index_e.php; location in figs 1, 3). the solid black line shows the profile predicted if all units in the well are currently at their maximum post-depositional temperatures (default thermal history, dth). all values plot consistently above the black line, suggesting that units below the plio-pleistocene saglek formation have been hotter in the past, although a detailed interpretation of these data is not possible because of the scatter in the data and differences between datasets. the red profile shows the prediction from a history in which the pre-saglek section has been buried more deeply by 750 m prior to exhumation. 9494 vertical movements both prior to and after break-up; on the island of newfoundland, late carboniferous and jurassic phases of exhumation removed kilometre-thick covers over palaeozoic basins (hendriks et al. 1993; williams et al. 1998), and in labrador, cretaceous sediments rest on precambrian basement (dorf 1967). in the offshore domain, several major hiatuses punctuate the stratigraphic record; e.g. along labrador between the palaeozoic sediments and precambrian basement and between lower cretaceous volcanics and underlying palaeozoic sediments, intra-cretaceous and base-tertiary unconformities and several intratertiary unconformities (mid-paleocene, late eocene, midoligocene and miocene; ainsworth et al. 2014). it is our ambition to combine the evidence from the stratigraphic record with results from stratigraphic landscape analysis and thermochronology to provide a coherent model of the timing and magnitude of the vertical movements along the margin both prior to and after break-up. acknowledgements we thank sponsoring oil companies for financial support and the geological survey of newfoundland and labrador for funding the afta pilot study. references ainsworth, n.r., riley, l., bailey, h.w. & gueinn, k.j. 2014: cretaceous–tertiary stratigraphy of the labrador shelf, riley geoscience ltd., commissioned by nalcor energy, http://www.nalcorenergy.com/ oilgas/labrador-biostratigraphy.asp avery, m.p. 2001: vitrinite reflectance (ro) of dispersed organic matter from husky/bow valley et al. golconda c-64. gsc open file report 4013, 14 pp., http://basin.gdr.nrcan.gc.ca/wells/single_maturation_e. php?well=d302 bonow, j.m., japsen, p., lidmar-bergström, k., chalmers, j.a. & pedersen, a.k. 2006: cenozoic uplift of nuussuaq and disko, west greenland – elevated erosion surfaces as uplift markers of a passive margin. geomorphology 80, 325–337. cooper, m., weissenberger, j., knight, i., hostad, d., gillespie, d., williams, h., burden, e., porter-chaudhry, j., rae, d. & clark, e. 2001: basin evolution in western newfoundland: new insights from hydrocarbon exploration. aapg bulletin 85, 393–418. dickie, k., keen, c.e., williams, g.l. & dehler, s.a. 2011: tectonostratigraphic evolution of the labrador margin, atlantic canada. marine and petroleum geology 28, 1663–1675. doré, a.g., cartwright, j.a., stoker, m.s., turner, j.p. & white, n. (eds) 2002: exhumation of the north atlantic margin: timing, mechanisms and implications for petroleum exploration. geological society, london, special publications 196, 494 pp. dorf, e. 1967: cretaceous insects from labrador i. geologic occurrence. psyche 74, 267–269. green, p.f., lidmar-bergström, k., japsen, p., bonow, j.m. & chalmers, j.a. 2013: stratigraphic landscape analysis, thermochronology and the episodic development of elevated passive continental margins. geological survey of denmark and greenland bulletin 30, 150 pp. grist, a. & zentilli, m. 2003: post-paleocene cooling in the southern canadian atlantic region: evidence from apatite fission track models. canadian journal of earth sciences 40, 1279–1297. hendriks, m., jamieson, r.a., willett, s.d. & zentilli, m. 1993: burial and exhumation of the long range inlier and its surroundings, western newfoundland: results of an apatite fission-track study. canadian journal of earth sciences 30, 1594–1606. japsen, p., bonow, j.m., green, p.f., chalmers, j.a. & lidmar-bergström, k. 2006: elevated, passive continental margins: long-term highs or neogene uplifts? new evidence from west greenland. earth and planetary science letters 248, 315–324. japsen, p., chalmers, j.a., green, p.f. & bonow, j.m. 2012: elevated, passive continental margins: not rift shoulders, but expressions of episodic, post-rift burial and exhumation. global and planetary change 90–91, 73–86. mcgregor, e.d., nielsen, s.b., stephenson, r., petersen, k.d. & macdonald, d.i.m. 2013: long-term exhumation of a palaeoproterozoic orogen and the role of pre-existing heterogeneous thermal crustal properties: a fission-track study of se baffin island. journal of the geological society, london 170, 877–891. mcmillan, n.j. 1973: shelves of labrador sea and baffin bay, canada. in: mccrossan, r.g. (ed.): future petroleum provinces of canada, their geology and potential. canadian society of petroleum geology, memoir 1, 473–517. pe-piper, g. & mackay, r.m. 2006: provenance of lower cretaceous sandstones onshore and offshore nova scotia from electron microprobe geochronology and chemical variation of detrital monazite. bulletin of canadian petroleum geology 54, 366–379. sinclair, i.k., shannon, p.m., williams, b.p.j., harker, s.d. & moore, j.g. 1994: tectonic control on sedimentary evolution of three north atlantic borderland mesozoic basins. basin research 6, 193–217. weissel, j.k. & karner, g.d. 1989: flexural uplift of rift f lanks due to mechanical unloading of the lithosphere during extension. journal of geophysical research – solid earth 94, 13919–13950. white, t.s., witzke, b.j. & ludvigson, g.a. 2000: evidence for an albian hudson arm connection between the cretaceous western interior seaway of north america and the labrador sea. geological society of america bulletin 112, 1342–1355. williams, s.h., burden, e.t. & mukhopadhyay, p.k. 1998: thermal maturity and burial history of paleozoic rocks in western newfoundland. canadian journal of earth sciences 35, 1307–1322. authors’ addresses p.j., geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: pj@geus.dk p.f.g., geotrack international, 37 melville road, brunswick west, victoria 3055, australia. j.m.b., geovisiona ab, högbyvägen 168, se-17554 järfälla, sweden and mid sweden university, kunskapens väg 1, se-831 25 östersund, sweden. a.m.h., geological survey, department of natural resources, government of newfoundland and labrador, p.o. box 8700, st. john’s, canada nl a1b 4j6. d.h.c.w., memorial university of newfoundland, p.o. box 4200, st. john’s, canada nl a1c 5s7. geological survey of denmark and greenland bulletin 13, 2007, 13-16 the central graben in the north sea is a mature petroleum province with upper jurassic – lowermost cretaceous marine shale of the kimmeridge clay formation and equivalents as the principal source rock, and upper cretaceous chalk as the main reservoirs. however, increasing oil prices and developments in drilling technologies have made deeper plays depending on older source rocks increasingly attractive. in recent years exploration activities have therefore also been directed towards deeper clastic plays where palaeozoic depo sits may act as petroleum source rocks. carboniferous coaly sections are the most obvious source rock candidates. the gas fields of the major gas province in the southern north sea and north-west europe are sourced from the thick upper carboniferous coal measures, which contain hundreds of coal seams (drozdzewski 1993; lokhorst 1998; gautier 2003). north of the gas province upper carboni-ferous coal-bearing strata occur onshore in northern england and in scotland, but offshore in the north sea area they have been removed by © geus, 2007. geological survey of denmark and greenland bulletin 13, 13–16. available at: www.geus.dk/publications/bull are carboniferous coals from the danish north sea oil-prone? henrik i. petersen and hans p. nytoft 13 fig. 1. a: simplified map showing carboniferous basins in the north sea area. the danish central graben is also shown (grey area). adb, anglodutch basin; fab, forth approaches basin; imfb, inner moray firth basin; mnh, mid netherlands high; mnsh, mid north sea high; mv, midland vally; ntb, northumberland/tweed basin; nwgb, north-west german basin; omfb, outer moray firth basin; rb, ruhr basin; wgg, witch ground graben. based on ziegler (1990), besly (1998) and bruce & stemmerik (2003). b: present-day distribution of lower carboniferous littoral/fluvial/deltaic and shallow marine/deltaic deposits in the southern north sea area. the positions of the gert-2, gert-3, p-1 and svane-1 wells are shown. modified from lokhorst (1998). c: sedimentological log of the coal-bearing interval of the gert-2 well. the coaly interval is underlain by fluvial sediments and overlain by fossiliferous marine mudstones of the so-called marine unit. slightly modified from petersen & nytoft (2007). erosion. however, lower carboniferous strata are present offshore and have been drilled in the witch ground graben and in the north-eastern part of the forth approaches basin (fig. 1a), where most of the lower carbon iferous sediments are assigned to the sandstone/shale-dominated tayport for mation and to the coal-bearing firth coal formation (bruce & stemmerik 2003). highly oil-prone lower carboniferous lacustrine oil shales occur onshore in the midland valley, scotland, but they have only been drilled by a single well off shore and seem not to be regionally distributed (parnell 1988). in the southern part of the norwegian and uk central graben and in the danish central graben a total of only nine wells have encountered lower carboniferous strata, and while they may have a widespread occurrence (fig. 1b; bruce & stemmerik 2003) their distribution is poorly constrained in this area. the nearly 6000 m deep svane-1/1a well (fig. 1b) in the tail end graben encountered gas and condensate at depths of 5400–5900 m, which based on carbon isotope values may have a carboniferous source (ohm et al. 2006). in the light of this the source rock potential of the lower carboniferous coals in the gert-2 well (fig. 1c) has recently been assessed (petersen & nytoft 2007). lower carboniferous strata in the danish central graben in the danish central graben, lower carboniferous strata were drilled by the gert-2, gert-3 and p-1 wells (fig. 1b). the depth to the lower carboniferous ranges from 3289 m in the p-1 well to 4840 m in the gert-2 well. whereas the p-1 well reached caledonian basement after penetrating about 67 m of carboniferous sediments, the gert-2 well drilled 192 m of carboniferous strata before drilling terminated at about 5000 m depth within the carboniferous. the drilled carbon iferous section in the gert-2 well is principally non-marine (fluvial unit) and contains a coaly interval at about 4890 m (fig. 1c) that constitutes a transition to marine shales and shoreface and tidally influenced sandstones of the marine unit (petersen & nytoft 2007). the coals overlie a fluvial finingupward succession and are overlain by fossiliferous marine shales (fig. 1c). the coals formed in peat-forming coastal plain mires as shown by high sulphur contents (average 5.3 wt%) and the presence of framboidal pyrite (petersen & nytoft 2007). high contents of vitrinite (65–82 vol.%), derived from degradation of higher land plant woody material, indicate waterlogged, oxygen-deficient conditions in the precursor mires. although the proportion of more oil-prone liptinite constituents is generally small (4–8 vol.%), the paralic peat-forming conditions may be favourable for the oil gene ration potential of the resulting coals (petersen 2006). this raises the question: are the coals encountered in the gert-2 well oilor gas-prone? source rock quality and hydrocarbon generation capacity the average tmax of the coals is 448°c, which corresponds to a vitrinite reflectance of ~0.95%ro indicating that the coals are at the threshold of, or slightly within, the so-called ‘effective oil window’ (in which efficient oil expulsion occurs; sykes 2001; petersen 2006). in addition the hydrogen index (hi) values of the coals are very close to their himax values. during initial maturation the hi of coals increases to a maximum value, which is considered to be a better estimate of the generation potential of coal (sykes & snowdon 2002; petersen 2006). thus, at first glance hi values from 171–219 mg hc/g toc may suggest some potential for liquid petroleum formation (fig. 2). the type of generated petroleum is, however, determined by the paraffinicity of the organic matter, i.e. the proportion and length of hydrogen-bearing carbon chains (aliphatic chains) in the kerogen structure. the ability to generate and expel typical waxy terrestrial crude oil requires the presence of long-chain aliphatics with more than ~20–25 carbon atoms (isaksen et al. 1998; killops et al. 1998). fourier transform infrared spectroscopy (ftir) of the gert-2 coals clearly reveals a response in the aliphatic stretching region, but the response can mainly be assigned to isolated ch2 compounds, which are of no importance to the liquid petroleum generation potential (petersen & nytoft 2006). quantification of the proportion of long-chain ali phatics in the kerogen structure of the gert-2 coals by comprehensive chemical treatment (so-called ruthenium tetroxide catalysed oxidation; see petersen & nytoft 2006, 2007) demonstrates a negligible or extremely low amount of ali phatic chains with more than 18 carbon atoms. the dominance of shorter aliphatic chains strongly indicates that the coals are gasand condensate-prone. 14 fig. 2. hydrogen index versus tmax plot of non-extracted and solvent extracted coal samples from the gert-2 well. the average hydrogen index of the non-extracted coals is also shown. carboniferous coals are inherently gas-prone the above results are in line with the findings of petersen & nytoft (2006), who showed that carboniferous coals in general contain very minor proportions of long-chain aliphatics in the range c19–35 and are therefore inherently poorly suited to generate oil. thus, for carbon iferous coals only an effective gas/con den sate window exists. of the total amount of aliphatic chains in the range c12–35, carbon iferous coals contain on average about 20% in the c19–35 range (fig. 3a). in contrast, jurassic coals from the søgne basin in the north sea contain about 26%, while cenozoic coals contain on average as much as 55% (fig. 3a). the significantly higher proportion of longchain ali-phatics in the youngest coals seems to be related to the high amount of organic de-trital groundmass (fig. 3b). the groundmass consists of detrital vitrinitic and liptinitic organic matter that can be positively correlated to the long-chain ali-phatics in the kerogen structure (petersen & nytoft 2006). the oil-proneness of the cenozoic coals thus seems to be related to the evolution of more diversified plant communities, in cluding the appearance of angio sperms in the late cretaceous. limited expulsion efficiency and implications for en hanced gas-proneness in agreement with the kerogen structure the generated hydrocarbons from the lower carboniferous gert-2 coals are dominated by shorter-chain aliphatics. these do not facilitate expulsion (isak sen et al. 1998), and the generated hydro carbons remain trapped in the coals. this is sustained by a pronounced diffe rence in the hi of the non-extracted coals and the hi of the extracted coals: upon extraction the hi is on average reduced by 30% (fig. 2; petersen & 15 fig. 3. a: the proportion (%) of c19–35 long-chain aliphatics of the total amount of c12–35 aliphatics in a number of carboniferous, jurassic and cenozoic coals. b: the proportion (vol.%) of groundmass composed of detrital vitrinite and liptinite in a number of carboniferous, jurassic and cenozoic coals. 16 nytoft 2007). hence, the measured hi values of the gert-2 coals are strongly influenced by the trapped petroleum in the coals. the limited (or lack of ) expulsion maintains a low saturate/aromatic hydrocarbon ratio of the trapped petroleum, which according to pepper & dodd (1995) is less thermally stable than expelled oil that is dominated by saturated (aliphatic) hydrocarbons. for source rocks with hi values below 300 mg hc/g toc, intra-source rock cracking of hydrocarbons commences from 115–145°c (pepper & dodd 1995). the average hi of the gert-2 coals is 193 mg hc/g toc, and the vitrinite reflectance values suggest burial temperatures of 124–132°c, implying that intra-source rock oil-to-gas cracking of the trapped hydrocarbons may enhance the gas-proneness of the coals. observation of pyrolytic carbon in the coals may provide direct evidence for gas generation (petersen & nytoft 2007). concluding remarks as is the case with other carboniferous coals, the drilled lower carboniferous coals encountered in the gert-2 well, located at the northern margin of the danish central graben, are gas-prone. the gas-proneness is inherited from the coaly organic matter, which generally contains a small amount of oil-prone kerogen due to the lack of long-chain aliphatics. limited expulsion efficiency maintains a low saturate/aromatic ratio of the generated and trapped petroleum. the thermally less stable petroleum mixture promotes intrasource rock oil-to-gas cracking of the trapped hydrocarbons in the coals, which enhances their gas-proneness. the thin coaly section, present in the gert-2 well, has no economic significance. however, provided that the lower carbon i ferous coaly section is regionally distributed and the section elsewhere is thicker with a larger number of coal seams and/or thick sections of coaly shale, it can potentially be a gas source for deep plays in the danish central graben and adjacent areas. this is supported by the encountered gas in the svane-1/1a well (ohm et al. 2006). acknowledgements the study was part of a larger project financially supported by the danish natural science research council (grant 21-04-0605). references besly, b.m. 1998: carboniferous. in: glennie, k.w. (ed.): petroleum geology of the north sea: basic concepts and recent advances, 104–136. oxford: blackwell science ltd. bruce, d.r.s. & stemmerik, l. 2003: carboniferous. in: evans, d., graham, c., armour, a. & bathurst, p. (eds): the millennium atlas: petroleum geology of the central and northern north sea, 83–89. bath: geological society of london. drozdzewski, g. 1993: the ruhr coal basin (germany): structural evolution of an autochthonous foreland basin. international journal of coal geology 23, 231–250. gautier, d.l. 2003: carboniferous-rotliegend total petroleum system description and assessment results summary. u.s. geological survey bulletin 2211, 24 pp. isaksen, g.h., curry, d.j., yeakel, j.d. & jenssen, a.i. 1998: controls on the oil and gas potential of humic coals. organic geochemistry 29, 23–44. killops, s.d., funnell, r.h., suggate, r.p., sykes, r., peters, k.e., walters, c., woolhouse, a.d., weston, r.j. & boudou, j.-p. 1998: predicting generation and expulsion of paraffinic oil from vitrinite-rich coals. organic geochemistry 29, 1–21. lokhorst, a. (ed.) 1998: nw european gas atlas – composition and isotope ratios of natural gases. utrecht: netherlands institute of applied geoscience (cd-rom). ohm, s.e., karlsen, d.a., roberts, a., johannessen, e. & høiland, o. 2006: the paleocene sandy siri fairway: an efficient ‘pipeline’ draining the prolific central graben? journal of petroleum geology 29, 53–82. parnell, j. 1988: lacustrine petroleum source rocks in the dinantian oil shale group, scotland: a review. in: fleet, a.j., kelts, k. & talbot, m.r. (eds): lacustrine petroleum source rocks. geological society special publication (london) 40, 235–246. pepper, a.s. & dodd, t.a. 1995: simple kinetic models of petroleum formation. part ii: oil-gas cracking. marine and petroleum geology 12, 321–340. petersen, h.i. 2006: the petroleum generation potential and effective oil window of humic coals related to coal composition and age. international journal of coal geology 67, 221–248. petersen, h.i. & nytoft, h.p. 2006: oil generation capacity of coals as a function of coal age and aliphatic structure. organic geochemistry 37, 558–583. petersen, h.i. & nytoft, h.p. 2007: assessment of the petroleum generation potential of lower carboniferous coals, north sea: evidence for inherently gas-prone source rocks. petroleum geoscience 13, 271–285. sykes, r. 2001: depositional and rank controls on the petroleum potential of coaly source rocks. in: hill, k.c. & bernecker, t. (eds): eastern australasian basins symposium, a refocused energy perspective for the future. petroleum exploration society of australia special publication 1, 591–601. sykes, r. & snowdon, l.r. 2002: guidelines for assessing the petroleum potential of coaly source rocks using rock-eval pyrolysis. organic geochemistry 33, 1441–1455. ziegler, p.a. 1990: geological atlas of western and central europe, 2nd edition, 239 pp. mijdrecht: shell international petroleum. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: hip@geus.dk geological survey of denmark and greenland bulletin 20, 2010, 35–38 35 the geological survey of denmark and greenland (geus) and the institute of baltic sea research in warnemünde (formerly the institut für meereskunde of the ddr) have co-operated for more than two decades on unravelling the history of the south-western part of the baltic sea, mainly based on shallow seismic profiling, sampling of sediment cores and analyses of core samples (jensen et al. 2002). here we report on some results from one of the latest joint cruises with the german research vessel maria s. merian. the baltic sea is one of the largest brackish-water seas in the world. however, during wide periods of the quaternary, the baltic sea area was either covered by the scandinavian ice sheet, or was a lake or a land area. well-dated marine deposits are only known from the last interglacial stage (the eemian) and from the holocene. during the eemian, connections to the baltic sea were found via karelia to the white sea, and via denmark and northern germany to the north sea (funder et al. 2002). during the holocene, a connection to the north sea was first established during the yoldia sea stage via south-central sweden, and later during the littorina sea stage via the danish/german/swedish straits (björck 1995). in addition to holocene and eemian deposits, preholocene marine sequences from germany, poland, estonia and latvia have been referred to the holsteinian, the saalian and the weichselian (e.g. kalnina 2001). in poland seven marine sequences of weichselian ages were reported by gałąazka & marks (2009). however, the chronology and correlation of these deposits are uncertain. during regional mapping in the early 1990s, the geological survey of sweden discovered pre-holocene marine sediments to the north-east of kriegers flak in the western part of the arkona basin (fig. 1a). this find has major palaeogeographical implications because it is the only interstadial marine deposit reported from the area. the occurrence was described by klingberg (1998). however, the age of the deposit was uncertain, with the only hint coming from a nonfinite radiocarbon age determination which showed that the deposit is older than 40  000 14c years (lab. no. ua-4116). recently, five samples from non-marine sediments deposited above the brackish unit at kriegers flak, but in connection with it, have been radiocarbon dated – and gave finite ages amino acid analysis of pre-holocene foraminifera from kriegers flak in the baltic sea ole bennike and bernd wagner denmark kriegers flak sweden germany arkona basin core 318190 b10 km k ar el ia white sea baltic sea arkona basin north sea jylland sweden finland norway denmark øresund a roar kriegers flak500 km fig. 1. a: map of north-western europe showing the location of kriegers flak and the location of place names mentioned in this paper. b: bathymetrical map of the kriegers flak area, showing the location of the studied core. the water depth varies from c. 15 m (the lightest areas) to c. 45 m (the darkest areas). © geus, 2010. geological survey of denmark and greenland bulletin 20, 35–38. open access: www.geus.dk/publications/bull 3636 of c. 36 000 – 41 000 calendar years bp, indicating a middle weichselian age (anjar et al. 2010). klingberg (1998) noted that the sequence consists of stiff clay underlain and overlain by till. part of the marine clay contains species-poor foraminiferal faunas dominated by elphidium excavatum and elphidium albiumbilicatum. the fauna implies brackish-water conditions, which was confirmed by extremely low δ18o values of carbonate from the foraminifera tests at between –11.2 and –11.9‰ (klingberg 1998). ordinary sea water and carbonates precipitated from it have values close to zero, whereas the greenland ice sheet ranges between –32 and –44‰ (north greenland ice core project members 2004). a reasonable way to explain the low values from kriegers flak is by assuming a mixture of sea water and meltwater from the scandinavian ice sheet. five samples from the clay were analysed for pollen, but the interpretation of the pollen spectra was hampered by the presence of reworked pollen grains (klingberg 1998). in order better to constrain the age of the deposit, new material was collected and a sample of foraminifera tests analysed for the ratio between l-isoleucine and d-alloisoleucine. since its development in the late 1960s, amino acid geochronology has been increasingly used for dating and correlating late cenozoic deposits that are beyond the range of radiocarbon dating (miller & brigham-grette 1989). most studies are performed on mollusc shells and tests of foraminifera. protein in live organisms consists of amino acid molecules in the l-isomer form. after the death of the organism, some of the amino acid molecules change to the d-form until an equilibrium is reached. the d/l ratio depends on the time elapsed since the death of the organism, the diagenetic temperature history and the species analysed. if the temperature history of a fossil sample is known, it is possible to calculate its age. this, however, is very difficult for the late quaternary which is known for its large and rapid temperature shifts. yet, sites with similar diagenetic temperature histories may be correlated even if that history is unknown. we suggest that deposits in the southern baltic have experienced approximately similar temperature histories as those in the north sea and will compare with sites of known age there. we do not consider onshore deposits because these may have experienced different diagenetic temperature histories. the pioneering study on amino acid geochronology in north-west europe by miller & mangerud (1985) used mollusc shells from onshore marine interglacial deposits. following a pilot study by sejrup et al. (1984), amino acid stratigraphy using foraminiferal tests has been widely used for dating and correlating interglacial marine sequences in the north sea region (e.g. knudsen & sejrup 1988; sejrup & knudsen 1993, 1999). in the present study, amino acid analysis was carried out to obtain more information about the age of the deposits from kriegers flak. for this purpose, we used tests of the foraminifera elphidium excavatum as this species has already been widely used in other studies of amino acid ratios in the region. assuming that kriegers flak has experienced a similar temperature history as denmark and the north sea, we only compare the d/l ratio of elphidium excavatum with d/l ratios from the same species, in order to avoid taxonomical effects which can be significant (e.g. miller & mangerud 1985). material and methods vibro-coring was carried out in the north-eastern swedish part of kriegers flak, at the position 55°04.08´n, 13°11.63´w, at a water depth of 39.7 m (fig. 1b). a 6 m long corer was used, but it only penetrated 315 cm. the core (no. 318190, institute of baltic sea research) consisted of diamicton (315–205 cm), light, olive-grey, stiff clay (205–34 cm), diamicton (34–10 cm) and silt with shells of marine molluscs (10–0 cm; fig. 2). a series of samples from the clay unit was dried and wet sieved on 0.4, 0.2 and 0.1 mm sieves. most of the samples were barren or only contained rare foraminifera, but one of the samples (from 198–193 cm core depth) contained abundant tests of the benthic foraminifera elphidium excavatum. the foraminifera tests were picked out and split into two subsamples, each weighing around 6 mg. both subsamples were analysed in the amino acid geochronology laboratory at the institute of arctic and alpine research, fig. 2. lithological log of the sediment core. d ep th b el o w c o re t o p ( cm ) holocene silt with shells stiff diamicton stiff clay stiff diamicton 0 40 80 120 160 200 240 280 37 boulder, colorado. three measurements were carried out on each subsample. peptide-bound amino acids were decomposed in the laboratory by hydrolysis, and the analyses were made with a chromatograph. the ratio of d-alloisoleucine to l-isoleucine (d/l) is based on the peak height in the total population, i.e. both free and peptide-bound amino acids. results subsamples aal-11579a and b gave d/l ratios of 0.106 ± 0.000 and 0.106 ± 0.002, respectively. although most analyses of foraminifers in north-west europe have been carried out at the amino acid chronological laboratory at the geological institute, bergen university, we regard the amino acid analyses from the institute of arctic and alpine research as nearly identical to those from bergen, since samples processed at both places have yielded similar results (miller & mangerud 1985). discussion of age amino acid analyses of numerous samples of elphidium excavatum tests from various interglacial and interstadial deposits from the north sea region have been summarised by sejrup & knudsen (1993, 1999). in the latter work, the authors divided the north sea sequence into four amino zones. zone 1 corresponds to the holocene and the late weichselian, zone 2 includes samples of eemian age, as well as samples of late saalian and early weichselian age, zone 3 includes samples of holsteinian age and zone 4 is correlated with marine isotope stage 11 (420–360 ka). the d/l ratios in amino zone 2 range from 0.08 to 0.12, and those in amino zone 3 show ratios between 0.14 and 0.16. subsamples aal-11579a and b from kriegers flak have ratios similar to those of amino zone 2, and thus imply an eemian, late saalian or early weichselian age. eemian age? an eemian age for the clay sequence from kriegers flak is unlikely, since the foraminiferal faunas consist almost entirely of the two benthic species elphidium excavatum and elphidium albiumbilicatum, an assemblage indicating brackish waters in an arctic/subarctic environment (klingberg 1998). the fauna is clearly different from eemian foraminiferal faunas from the region, as they are dominated by species-rich assemblages implying warmer conditions. elphidium excavatum is common in late eemian deposits at mommark on als, southern denmark (e.g. kristensen & knudsen 2006). however, the eemian deposits at this site show normal marine oxygen isotope ratios, whereas the ratios at kriegers flak seem to suggest deposition during a time with extensive melting of glacial ice as noted above. early weichselian age? an early weichselian age is also considered unlikely. during the early weichselian, the global sea level was lower than at present (e.g. siddall et al. 2003) and the north sea as well as the danish/german/swedish straits are generally believed to have been dry land (houmark-nielsen 1989). according to the glaciation curves of e.g. lundquist (1986), houmark-nielsen (1989) and mangerud (2004), the scandinavian ice sheet did not advance to the south-western baltic during the early weichselian. hence, arctic/subarctic brackish-water conditions with marked meltwater influence would not be expected in the kriegers flak region during this time interval. late saalian age? during the late saalian, the sea transgressed large parts of the north sea and reached as far south as the roar, where boreo-arctic and arctic conditions were found (fig. 1a; knudsen 1986). in northern denmark, the børglum chronozone of lykke-andersen & knudsen (1991) reflects interstadial conditions of late saalian age. deposits from an arctic sea have also been recorded in southern sweden (påsse et al. 1988). according to the reconstruction by houmark-nielsen (1989), the margin of the scandinavian ice sheet retreated from the south-western baltic sea. we consider it possible that the øresund region was isostatically depressed following the extensive late saalian glaciation and transgressed by the sea for a short time interval after recession of the ice. brackish waters may have extended to the south-western part of the baltic sea, including kriegers flak. middle weichselian age? a late saalian age is obviously not in accordance with the radiocarbon dates reported by anjar et al. (2010). the radiocarbon dates are close to the limit of radiocarbon dating and it could be argued that they are unreliable, in particular because four of them were made on bulk sediment samples. dating of bulk sediment samples commonly yields erroneous results, but usually dates of bulk samples are too old – not too young. also, all five age determinations gave finite dates, which is a good indication that they are reliable. we find it difficult to reconcile a middle weichselian age with the amino acid data, which suggest an older age. however, in the past decade it has become more and more obvious that amino acid ratios should be used with caution due to their strong dependency on the temperature histories of the deposits. hence, a middle weichselian age cannot be excluded from the amino acid data. figure 3 shows a possible palaeogeographical reconstruction. we suggest that the øresund region was isostatically depressed and transgressed by the sea for a short time interval after recession of the scandinavian ice sheet. 3838 conclusions the amino acid ratios from brackish-water deposits on kriegers flak suggest a late quaternary age, either late saalian, eemian or early weichselian. a strong δ18o meltwater signal and the foraminiferal assemblage are not compatible with an eemian age. an early weichselian age is also considered unlikely because the sea level was probably too low at this time to allow the sea to transgress the baltic sea, and because the margin of the scandinavian ice sheet was too far away to give a strong meltwater signal. the last option, a late saalian age, does not agree with a middle weichselian age of the fresh-water deposits that are connected to the brackish-water deposits. hopefully, other dating methods can be applied at kriegers flak in the future to better constrain the age of the sequence. references anjar, j., larsen, n.k., björck, s., adrielsson, l. & filipsson, h.l. 2010: mis 3 marine and lacustrine sediments at kriegers flak, southwestern baltic sea. boreas 39, 360–366. björck, s. 1995: a review of the history of the baltic sea, 13-8 ka bp. quaternary international 27, 19–40. funder, s., demidov, i. & yelovicheva, y. 2002: hydrography and mollusc faunas of the baltic and the white sea – north sea seaway in the eemian. palaeogeography, palaeoclimatology, palaeoecology 184, 275–304. gałąazka, d. & marks, l. 2009: geology of the lower vistula region, northern poland. polish geological institute, special papers 25, 13–20. houmark-nielsen, m. 1989: the last interglacial–glacial cycle in denmark. quaternary international 3/4, 31–39. jensen, j.b., kuijpers, a., bennike, o. & lemke, w. 2002: balkat. the baltic without frontiers. geologi – nyt fra geus 4, 19 pp. copenhagen: geological survey of denmark and greenland. kalnina, l. 2001: middle and late pleistocene environmental changes recorded in the latvian part of the baltic sea basin. quaternaria a9, 173 pp. klingberg, f. 1998: a late pleistocene marine clay succession at kriegers flak, westernmost baltic, southern scandinavia. journal of quaternary science 13, 245–253. knudsen, k.l. 1986: middle and late quaternary foraminiferal stratigraphy in the southern and central north sea area. striae 24, 201–205. knudsen, k.l. & sejrup, h.p. 1988: amino acid geochronology of selected interglacial sites in the north sea area. boreas 17, 347–354. kristensen, p.h. & knudsen, k.l. 2006: palaeoenvironments of a complete eemian sequence at mommark, south denmark: foraminifera, ostracods and stable isotopes. boreas 35, 349–366. lundquist, j. 1986: stratigraphy of the central area of the scandinavian glaciation. quaternary science reviews 5, 251–268. lykke-andersen, a.l. & knudsen, k.l. 1991: saalian, eemian and weichselian in the vendsyssel–kattegat region, denmark. striae 34, 135–140. mangerud, j. 2004: ice sheet limits on norway and the norwegian continental shelf. in: ehlers, j. & gibbard, p. (eds): quaternary glaciations – extent and chronology 1, 271–294. amsterdam: elsevier. miller, g.h. & brigham-grette, j. 1989: amino acid geochronology: resolution and precision in carbonate fossils. quaternary international 1, 111–128. miller, g.h. & mangerud, j. 1985: aminostratigraphy of european marine interglacial deposits. quaternary science reviews 4, 215–278. north greenland ice core project members 2004: high-resolution record of northern hemisphere climate extending into the last interglacial period. nature 431, 147–151. påsse, p., robertsson, a.-m., miller, u. & klingberg, f. 1988: a late pleistocene sequence at margreteberg, southwestern sweden. boreas 17, 141–163. sejrup, h.p. & knudsen, k.l. 1993: paleoenvironments and correlations of interglacial sediments in the north sea. boreas 22, 223–235. sejrup, h.p. & knudsen, k.l. 1999: geochronology and palaeoenvironment of marine quaternary deposits in denmark: new evidence from northern jutland. geological magazine 136, 561–578. sejrup, h.p., rokoengen, k. & miller, g.h. 1984: isoleucine epimerization in quaternary benthonic foraminifera from the norwegian continental shelf: a pilot study. marine geology 56, 227–239. siddall, m., rohling, e.j., almogi-laban, a., hemleben, c., meiscner, d., schmelzer, i. & smeed, d.a. 2003: sea-level fluctuations during the last glacial cycle. nature 423, 853–858. scandinavian ice sheet kriegers flak core 250 km fig. 3. palaeogeographical model of southern scandinavia during deposition of the kriegers flak sequence. authors’ addresses o.b., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: obe@geus.dk b.w., baltic sea research institute, seestrasse 15, d-18119 rostock-warnemünde, germany. present address: institute for geology and mineralogy, university of cologne, zülpicher str. 49a, d-50674 cologne, germany. geological survey of denmark and greenland bulletin 20, 2010, 91–94 91 a number of sedimentary basins of various ages are located on and offshore vietnam (fig. 1). some of them have significant petroleum resources and have thus attracted interest from industry and academia (rangin et al. 1995; matthews et al. 1997; lee & watkins 1998; lee et al. 2001). moreover, vietnam is located in a position central to the understanding of the geological development of south-east asia (hall & morley 2004). the structural style and the stratigraphy of the vietnamese basins thus provide a valuable record about the development of south-east asia throughout the phanerozoic and the subsequent eocene as well as younger deformation associated with the collision and indentation of india into eurasia and the opening of the south china sea (fyhn et al. 2009a, 2010a). the geological survey of denmark and greenland has worked in vietnam since 1995 to assess the geology and petroleum potential of the vietnamese basins. since 2002 the work has been carried out in cooperation with the department of geography and geology, university of copenhagen, as part of the enreca project (enhancement of research capacity in developing countries). an integrated part of the project is its training of vietnamese msc and phd students incorporating both training courses at the department of geography and geology and courses held in vietnam. so far, 10 msc and 4 phd students have completed their training under the auspices of the enreca project and another 10 are expected to complete their education within the next phase of the project. the enreca project has already completed two phases and a third and final phase has recently started. the initial phase focused on the phu khanh and the song hong basins located in the south china sea offshore north and central vietnam and the smaller onshore song ba trough (fig. 1; bojesen-koefoed et al. 2005; nielsen et al. 2007; fyhn et al. 2009a, b, c). during the second enreca phase, completed in 2009, attention shifted towards the malay – tho chu and phu quoc basins located in the gulf of thailand, ssw of vietnam (petersen et al. 2009, in press; fyhn et al. 2010a, b). the phu quoc basin continues onshore to the north to form part of the mountainous area between vietnam and cambodia. in the recently started third phase of the project, the focus remains on the phu quoc basin in addition to a revisit to the song hong basin on the north vietnamese margin and onshore beneath the song hong (red river) delta. the phu quoc basin the phu quoc basin stretches in a 100–150 km broad belt from the central part of the gulf of thailand c. 500 km northwards to central cambodia. the basin is late jurassic to cretaceous in age but is one of the least explored basins in the region and remains to be drilled offshore (fyhn et al. 2010a). in order to assess the geological evolution and the petroleum potential of the basin, regional seismic analyses of the vietnamese part of the basin were carried out in combination with drilling of the fully cored, 500 m deep enreca-2 well on the phu quoc island. data from the enreca-2 well were complemented by data from outcrop studies on phu quoc and in cambodia. vietnamese sedimentary basins: geological evolution and petroleum potential michael b.w. fyhn, henrik i. petersen, anders mathiesen, lars h. nielsen, stig a.s. pedersen, sofie lindström, jørgen a. bojesen-koefoed, ioannis abatzis and lars o. boldreel phu khanh basin phu khanh basingulf of thailand gulf of thailand � � � � � � � � � �� red river shear zone ? m ae ping shear zone w. natuna basin w. natuna basin penyu basin mtcbmtcb mergui basin mergui basin pattani basin pattani basin malay basin malay basin fig.2 andaman sea three pagodas shear zone malaysia cambodia vietnam laos myanmar thailand sedimentary basin major thrust fault oceanic crust shear zone with main cenozoic strike-slip direction 100°e 110°e 15°n sagain fau lt w . a n d am an fault ? ? ? nam con son basin nam con son basin cuu long basin cuu long basin phu quoc basin song hong basin song hong basin ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ india south china sea china borneo australia ▲ ▲ ▲ ▲ ▲ ▲ ▲ 2000 km khorat basin khorat basin song ba trough song ba trough phu quoc basin 250 km phu quoc basin © geus, 2010. geological survey of denmark and greenland bulletin 20, 91–94. open access: www.geus.dk/publications/bull fig. 1. map of south-east asia showing the locations of sedimentary basins and areas underlain by oceanic crust. strike-slip arrows illustrate the prevailing eocene–recent offset directions. mtcb: malay – tho chu basin. the inset map shows a simplified structural outline of the region. modified from fyhn et al. (2010b). 9292 alluvial sandstones with an average of c. 10% rhyolitedominated lithic fragments make up the greater part of the up to c. 4 km thick sediments filling the phu quoc basin. only a few, minor, shallow marine sandstone beds have been encountered in the terrestrially dominated succession. the sandstone-dominated succession is intercalated with subordinate alluvial plain and lacustrine siltand mudstone intervals. coal fragments are abundant at specific stratigraphic levels but do not contribute to any source potential. the thicknesses of the deposits are not affected by synsedimentary faulting, but gradually increase towards the east, where a coeval, jurassic–cretaceous, magmatic arc parallels the eastern basin flank (fig. 2). this is compatible with a retroarc–foreland basin setting associated with the growth of the magmatic arc located east of the basin that also served as a primary source of sediments for the basin. a distinct basin inversion is indicated by a prominent angular unconformity that caps the mesozoic basin fill and is associated with spectacular thrust faulting and folding (fig. 3). the structural complexity increases towards the deeply eroded and hitherto undescribed fold belts that confine the basin to the east and west. the stratigraphic level of erosion increases towards these orogenic belts. palaeozoic and lower mesozoic igneous and sedimentary rocks therefore crop out on small islands and onshore, or subcrop towards the base of the cenozoic within the kampot fold belt flanking the basin to the east. the inversion unconformity is underlain by lower cretaceous deposits and overlain by middle eocene and younger deposits, which provide only modest information on the age of the orogenic event. in order to constrain the age of inversion more precisely, apatite fission track analysis (afta) was carried out on rock samples from the kampot fold belt collected on islands and in mainland vietnam. the afta samples demonstrate a distinct cooling event that affected the region during the period from late paleocene to early eocene (fyhn et al. 2010a). the cooling corresponds to uplift and denudation of the area in response to the thrust faulting and basin inversion and thus confines the age of the orogenic event controlling the deformation. the malay – tho chu basin the malay – tho chu basin constitutes the vietnamese north-eastern part of the malay basin that was initiated somewhere between the middle and late eocene (fyhn et al. 2010b). the malay – tho chu basin is situated in the central part of the gulf of thailand and is therefore superimposed on the southernmost part of the phu quoc basin. rifting in the area took place between middle?–late eocene and oligocene time and resulted in the creation of a series of deep grabens filled by continental to shallow marine deposits (fig. 4; fyhn et al. 2010b). a set of nnw-trending rifts are the dominating structures in the area; they are often distinguished from other, wnw-trending rifts by their downwards steepening, their great depth and their linearity. the nnw-trending rifts seem to have accommodated leftlateral transtension. large-scale eocene–oligocene rifting associated with strike-slip faulting in the region is viewed as a response to the indian–eurasian collision that forced the neighbouring parts of south-east asia away from the collision zone through a series of lateral shear zones (e.g. tapponnier et al. 1986; fyhn et al. 2009a, 2010b), although n k h m e r f o ld b e lt k a m p o t f o ld b e lt ju r a s s ic – c re ta c e o u s m a g m a tic a rc u. jurassic – l. cretaceous strongly deformed u. jurassic – l. cretaceous pre-u. jurassic mainly mesozoic igneous rocks middle to upper cenozoic extensional fault subcrop concealed under thick cenozoic cover lower cenozoic contractional fault 50 km phu quoc enreca-2 vietnam m a lay – t ho c hu bas in cambodia phu quoc basin fig. 3 fig. 4 10°s 104°e fig. 2. subcrop map towards the base cenozoic top / mesozoic unconformity outlining the southern part of the phu quoc basin. fold belts confine the basin to the east and west. onshore pre-quaternary outcrops are indicated, outlining the onshore continuation of the phu quoc basin, the kampot fold belt and the se indochina jurassic–cretaceous magmatic arc. for location see fig. 1. modified from fyhn et al. (2010a). 93 other theories exist (morley 2002; hall & morley 2004; watkinson et al. 2008; hall 2009). around the onset of the miocene, rifting declined and thermal sagging came to dominate throughout the neogene. this resulted in broadening of the subsiding area and increasing marine influence. deltaic and shallow marine siliciclastics therefore prevail in the neogene succession of the basin. petroleum exploration in the malay – tho chu basin began during the early 1970s, encouraged by successful exploration immediately south of vietnamese territory. the first well was drilled in 1994. since then, significant gas, condensate and oil discoveries have been made in several wells drilled in this basin, but only a few discoveries are as yet considered commercial. a re-evaluation of the existing exploration strategies is therefore desirable in order to optimise and focus future exploration. exploration has mainly aimed at early to middle miocene fluviodeltaic sand reservoirs with late neogene structural trapping mechanisms. potential source rocks have been interpreted as alginite-bearing lacustrine shales and humic coals situated in the palaeogene synrift and in the lowermost post-rift sections. the few potential source-rock levels penetrated by wells have shown remarkably low vitrinite reflectance (vr) values compared to vr values obtained from overlying neogene coals (data presented in petersen et al. 2009). the maturity trends of such vr data sets are not well constrained as they produce suspiciously low thermal maturity gradients. however, suppressed vr values may occur in algenite-rich rocks. vr suppression is therefore particularly common in lacustrine shales with high hi (hydrogen index) values. the fluorescence alteration of multiple macerals (famm) technique is an accurate method to determine thermal maturity in rocks including those containing vitrinite with suppressed and enhanced vr values (willkins et al. 1992). by combining conventional vr measurements and the famm technique a revised, higher and more reliable thermal maturity gradient has been established (petersen et al. 2009). 2-d modelling of the maturation history of the basin was carried out based on the revised thermal maturity gradient, detailed seismic mapping, borehole information and new custom kinetics for petroleum generation; the latter was determined from lacustrine source rock samples and a terrestrially influenced mudstone collected from wells (petersen et al. in press). maturation modelling indicates that most of the syn-rift succession in the vietnamese malay basin is located in, or has passed through, the main oil and gas windows. carbonaceous 0 1 2 3 5 km tw o -w ay t ra ve l ti m e (s ec ) 4 5 sw l. pliocene u. miocene m. miocene pr ec en oz oi c eo ce ne l. miocene u. oligocene ne wnw-trending normal fault wnw-trending normal fault nnw-trending transtensional faults dhisdhi fig. 4. seismic transect across a nnwtrending palaeogene graben bounded by steep transtensional faults and half grabens confined by more gently dipping wnw-trending normal faults that link up with the steep faults at depth. deposition broadened across basement highs following the palaeogene synrift period due to regional thermal sagging. dhis (direct hydrocarbon indicators) within the miocene succession are frequently associated with structural traps formed during the late neogene. modified from fyhn et al. (2010b). 0 1 2 inversion unconformity ? sea neogene u. jurassic – l. cretaceous pre-jurassic tw o -w ay t ra ve l ti m e (s ec ) 10 km w e fig. 3. offshore stratigraphic profile across the phu quoc basin. jurassic–cretaceous subsidence was most intense simultaneous with the magmatic arc developing to the east. a prominent inversion unconformity separates the mesozoic thrust-faulted strata from the truncating unconformity of the base of the neogene deposits. modified from fyhn et al. (2010a). 9494 syn-rift deposits have therefore undergone adequate maturation and may have produced and expelled significant quantities of hydrocarbons. the oldest deposits in the deepest part of syn-rift depressions entered the oil window during the palaeogene syn-rift period but the main oil generation generally took place during early and middle miocene times. 2-d modelling of the hydrocarbon generation therefore indicates that the main risks in the tested play types are (1) the timing of petroleum generation relative to trap formation completed in the late neogene, (2) the pervasive faulting, which may have complicated petroleum migration to the structures and breached charged traps and (3) the distribution and amount of matured source rocks in smaller grabens. based on the above-mentioned criteria and the presence of direct hydrocarbon indicators (dhi), an untested alternative play type is proposed relying on syn-rift sandstones located up-dip from and near source-rock intervals with palaeogene structural and stratigraphic trapping mechanisms that did not experience subsequent neogene deformation. acknowledgements the enreca project is funded by the danish ministry of foreign affairs through danida. geocenter denmark provided additional financial support. vietnam petroleum institute (petrovietnam) is thanked for providing data and permission to publish this paper. references bojesen-koefoed, j.a., nielsen, l.h., nytoft, h.p., petersen, h.i., dau, n.t., hien, l.v., duc, n.a. & quy, n.h. 2005: geochemical characteristics of oil seepages from dam thi nai, central vietnam: implications for hydrocarbon exploration in the offshore phu khanh basin. journal of petroleum geology 28, 3–18. fyhn, m.b.w., boldreel, l.o. & nielsen, l.h. 2009a: geological development of the central and south vietnamese margin: implications for the establishment of the south china sea, indochinese escape tectonics and cenozoic volcanism. tectonophysics 478, 184–204. fyhn, m.b.w., boldreel, l.o. & nielsen, l.h. 2009b: tectonic and climatic control on growth and demise of the phanh rang carbonate platform offshore south vietnam. basin research 21, 225–251. fyhn, m.b.w. et al. 2009c: geological evolution, regional perspectives and hydrocarbon potential of the northwest phu khanh basin, offshore central vietnam. marine and petroleum geology 26, 1–24. fyhn, m.b.w., pedersen, s.a.s., boldreel, l.o., nielsen, l.h., green, p.f., dien, p.t., huyen, l.t. & frei, d. 2010a: palaeocene – early eocene inversion of the phuquoc – kampot som basin: se asian deformation associated with the suturing of luconia. journal of the geological society of london 167, 281–295. fyhn, m.b.w., boldreel, l.o. & nielsen, l.h. 2010b: escape tectonism in the gulf of thailand: paleogene left-lateral pull-apart rifting in the vietnamese part of the malay basin. tectonophysics 483, 365–376. hall, r. 2009: hydrocarbon basins in se asia: understanding why they are there. petroleum geoscience 15, 131–146. hall, r. & morley, c.k. 2004: sundaland basins. in: clift, p. et al. (eds): continent–ocean interactions within east asian marginal seas. geophysical monograph 149, 55–85. lee, g.h. & watkins, j.s. 1998: seismic sequence stratigraphy and hydrocarbon potential of the phu khan basin, offshore central vietnam, south china sea. aapg bulletin 82, 1711–1735. tulsa, oklahoma: american association of petroleum geologists. lee, g.h., lee, k. & watkins, j.s. 2001: geological evolution of the cuu long and nam con son basins, offshore southern vietnam, south china sea. aapg bulletin 85, 1055–1082. tulsa, oklahoma: american association of petroleum geologists. matthews, s.j., fraser, a.j., lowe, s., todd, s.p. & peel, f.j. 1997: structure, stratigraphy and petroleum geology of the se nam con son basin, offshore vietnam. in: fraser, a.j., matthews, s.j. & murphy, r.w. (eds): petroleum geology of southeast asia. geological society special publications (london) 126, 89–106. morley, c.k. 2002: a tectonic model for the tertiary evolution of strikeslip faults and rift basins in se asia. tectonophysics 347, 189–215. nielsen, l.h., petersen, h.i., thai, n.d., duc, n.a., fyhn, m.b.w., boldreel, l.o., tuan, h.a., lindstrøm, s. & hien, l.v. 2007: a middle– upper miocene fluvial–lacustrine rift sequence in the song ba rift, vietnam: an analogue to oil-prone, small-scale continental rift basins. petroleum geoscience 13, 145–168. petersen, h.i., sherwood, n., mathiesen, a., fyhn, m.b.w., dau, n.t., russell, n., bojesen-koefoed, j.a. & nielsen, l.h. 2009: application of integrated vitrinite reflectance and famm analyses for thermal maturity assessment of the northeastern malay basin, offshore vietnam: implications for petroleum prospectivity evaluation. marine and petroleum geology 26, 319–332. petersen, h.i., mathiesen, a., fyhn, m.b.w., dau, n.t., bojesen-koefoed, j.a., nielsen, l.h. & nytoft, h.p. in press: modeling of petroleum generation in the vietnamese part of the malay basin using custom kinetics. aapg bulletin. tulsa, oklahoma: american association of petroleum geologists. rangin, c., klein, m., roques, d., le pichon, x. & trong, l.v. 1995: the red river fault system in the tonkin gulf, vietnam. tectonophysics 243, 209–222. tapponnier, p., peltzer, g. armijo, r. 1986: on the mechanics of the collision between india and asia. in: coward, m.p. & ries, a.c. (eds): collision tectonics. geological society special publications (london) 19, 115–157. watkinson, i., elders, c. & hall, r. 2008: the kinematic history of the khlong marui and ranong faults, southern thailand. journal of structural geology 30, 1554–1571. wilkins, r.w.t., wilmshurst, j.r., russell, n.j., hladky, g., ellacott, m.v. & buckingham, c.p. 1992: fluorescence alteration and the suppression of vitrinite reflectance. organic geochemistry 18, 629–640. authors’ addresses m.b.w.f., h.i.p., a.m., l.h.n., s.a.s.p., s.l., j.a.b-k. & i.a., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: mbwf@geus.dk l.o.b., department of geography and geology, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. geological survey of denmark and greenland bulletin 17, 2009, 17-20 injection of co2 is a method that may increase the recovery of oil from danish chalk reservoirs in the north sea. the method is used elsewhere, particularly in north america, but has so far not been used in the north sea and has nowhere been used for chalk reservoirs, and the performance of the method when used for north sea chalk is therefore uncertain. a laboratory flooding experiment was conducted at the geological survey of denmark and greenland on a sample from the nana-1x well of the halfdan oil field in the danish north sea in order to test the efficiency of co2-enriched water to produce additional oil from chalk. the sample is a low-permeability chalk from the ekofisk formation and represents rocks that are marginal to the halfdan reservoir in an economical sense. outline of the experiment for the flooding experiment, four 1.5 inch core plug samples were assembled to form a composite sample with a total length of 28 cm and a pore volume of 92 ml. the flooding experiment was conducted at a fluid pressure of 282 bars, a hydrostatic confining pressure of 429 bars, and a temperature of 85°c. these are conditions similar to those of the halfdan reservoir. first the oil content, so, of the sample was adjusted to 77.7% of the pore volume (pv), the remaining pore fluid being simulated formation water. the sample was then brought to reservoir conditions, aged for three weeks to restore the wettability to reservoir conditions and then flood ed with simulated formation water until oil production from the sample had ceased. after changing the flooding fluid to co2-enriched water, flooding was resumed and sustained until oil production had declined to a negligible level. flood ing was then stopped, the rig was cooled and depressurised, and the sample was dismounted. both flooding operations were conducted with the sample in a vertical position from the bottom towards the top. a thorough description of the experiment is given in olsen (2007). experimental set-up the experiment was conducted in a rig that simulates reservoir conditions. the rig consists of a hassler-type core holder, a number of pressure cylinders for the experimental fluids, an acoustic separator for quantifying the fluid production, a differential pressure transducer for permeability measurement, and a high-pressure pump system for generating confining pressure, flow and pore fluid pressure. the core holder, pressure cylinders, separator, and differential pressure transducer are all situated inside a thermostatically controlled oven (fig. 1). temperature and fluid pressure conditions were based on data from the adjacent dan field reservoir, and corrected for the 250 m depth difference between the two reservoirs, half dan being the deeper. the temperature was corrected using a temperature gradient of 0.04°c/m. assuming pressure correspondence between the two reservoirs, the halfdan fluid pressure was estimated by extrapolation from the dan field using a pressure gradient of 0.075 bar/m. temperature, pressure and gradient values are from jørgensen (1992). the oil used in the experiment was degassed crude oil from the dan field. the water composition was similar to that of formation water from the halfdan field. fluid densities were measured at geus, while the water viscosity at reservoir conditions was estimated from data on the viscosity of similar brines. during the experiment differential pressure across the sample, pore fluid pressure, hydrostatic confining pressure, flow increased oil recovery from halfdan chalk by flooding with co2-enriched water: a laboratory experiment dan olsen © geus, 2009. geological survey of denmark and greenland bulletin 17, 17–20. available at: www.geus.dk/publications/bull 17 fig. 1. oven with high-pressure equipment used for the experiment described in this paper. the oven is 160 cm high. the arrow shows the core holder that was used for the experiment. rosa_2008:rosa-2008 01/07/09 15:47 side 17 rate, cumulative injected fluid volume, produced oil volume and temperature were continuously logged. before and after the flooding experiment the sample was characterised by measuring a number of parameters (table 1). water flooding the water flooding took place with a constant flooding rate of 0.62 ml/h and lasted 33 days with a total water injection of 490 ml or 5.35 times the pore volume. results of the water flooding are presented in fig. 2. water breakthrough oc curred after 77 hours when the water throughput was 0.458 times the pore volume. before breakthrough, oil was produced from the sample at the same rate as the water was being injected. after breakthrough, the rate of oil production drop ped sharply and continuously. a low oil production rate was sustained for a considerable time, but stopped completely before the flooding was terminated. a total oil volume of 0.053 times the pore volume was produced after breakthrough. total oil production during the water flooding was 0.496 times the pore volume or 63.8% of the oil originally present in the sample. a model was developed that fits the oil production (fig. 2). before breakthrough, the oil production shows a linear relationship with the injected water volume. after breakthrough, the oil production shows an exponentially decreasing relationship. both relationships show a nearly perfect fit to the actual oil production. measurements of differential pressure across the sample were used to calculate the water permeability during the water flooding (fig. 2). at the end of the water flooding the differential pressure had stabilised, indicating that fluid movement within the sample had stopped. both the oil production and permeability curves show typical water-flooding development. flooding with co2-enriched water the flooding with co2-enriched water was carried out at the same rate as the water flooding, i.e. at 0.62 ml/h. it lasted 64 days, and the total throughput of co2-enriched water was 949 ml or 10.4 times the pore volume. the co2-enriched water had a co2-content of 26.6 standard m3 co2/standard m3 water corresponding to a co2-saturation of 100% at 85°c and 282 bars fluid pressure (chang et al. 1998). diffusion of co2 between oil, co2-enriched water and water without co2 may cause the oil within the separator to either swell or shrink as co2 diffuses between the fluid phases. such volume changes are troublesome as they cannot be distinguished from oil being produced from the sample. in an attempt to establish equilibrium between separator oil and co2-enriched water, an amount of co2 was added to the separator before starting the co2-enriched flooding and allowed to equilibrate with the separator fluids for nine days. using the data of chang et al. (1998) the amount of co2 was adjusted so as to create the same co2-saturation in the water of the separator as in the brine used for flooding. figure 3 presents a plot of oil produced during the co2enriched flooding versus time. the oil production curve has 18 table 1. flooding experiment to enhance oil recovery: sample characterisation initial final change percent characterisation characterisation change dry weight (g) 613.37 609.05 –4.32 –0.70 diameter (cm) 3.77 3.77 0.00 0.07 length (cm) 28.32 28.24 –0.07 –0.26 bulk volume (ml) 318.56 316.92 –1.64 –0.52 porosity (% bulk volume) 28.77 29.09 0.32 1.10 pore volume (ml) 91.64 92.18 0.53 0.58 gas permeability (md) 0.62 0.76 0.14 23 produced oil model of produced oil permeability to waterpr od uc ed o il (f ra ct io n of p or e vo lu m e) 0 0.10 0.20 0.30 0.40 0.50 0 1 2 3 4 5 injected water volume (fraction of pore volume) 0.000 0.010 0.020 0.030 0.040 0.050 pe rm ea bi lit y to w at er ( m ill id ar cy ) fig. 2. produced oil volume (np) versus injected water volume (vinj) during flooding with water. model before breakthrough: np = 0.98 × vinj. model after breakthrough: np = 0.50 – 0.08 × exp(–vinj/1.57). fluid saturation (fraction of pore volume): initial water saturation = 0.223, water breakthrough at water saturation = 0.666, final water saturation = 0.719, produced oil = 0.496. rosa_2008:rosa-2008 01/07/09 15:47 side 18 a peculiar shape with an initial negative slope, indicating that the volume of oil in the separator was reduced. as oil cannot flow from the separator, the negative slope indicates that the oil in the separator shrank in volume during the first part of the co2-enriched flooding. after the section with negative slope, the slope of the oil production curve changes to positive and obtains the appearance of an ordinary oil production curve with oil apparently being produced at a low rate right until the end of flooding. after cessation of the flooding the rig was left undisturbed for 10.8 days, with the conditions of the rig being the same as during the co2-enriched flooding, except that the fluid delivery pump was stopped. during this time the separator continued to register an increase in oil volume, at a rate that was indistinguishable from the rate during the final part of the co2-enriched flooding (fig. 3). as the construction of the rig prevents oil from flowing to the separator when the delivery pump is stopped, the apparent oil production after flow-stop is considered to represent swelling of the oil within the separator. the situation appears similar to that at the beginning of the co2-enriched flooding, only that swelling takes place instead of shrinkage. during flooding with co2-enriched water it is expected that no oil is produced from the sample before oil affected by the co2 has moved to the outlet end of the sample. it is therefore reasonable to assume that the initial section of the oil production curve with negative slope represents the time before breakthrough of the co2-enriched water and the first produced oil. the section with negative slope then represents a period without oil production and should be horizontal. the section of the oil production curve after flow-stop should also be horizontal as no oil can be produced without flow. using these arguments a swelling correction curve has been constructed (fig. 3). as oil shrinkage in the separator changed to oil swelling during the experiment, a section with neither shrinkage nor swelling is present in the middle part of the correction curve. this part of the curve indicates equilibrium in the separator. the correction curve of fig. 3 therefore consists of three linear segments indicating shrinkage, equilibrium and swelling. a true correction curve probably would have a gradually changing slope, but in the absence of more information, linear line segments have been used. the data are interpreted as follows: at the start of flooding with co2-enriched water, the pore space of the sample contained 66 ml of co2-free water while the fluids of the separator were co2-saturated. during the first week of the flooding, the co2-free water of the sample flowed to the separator and caused the oil of the separator to shrink as co2 diffused from oil to water. after breakthrough of co2enriched water, the co2 content of the water in the separator gradually increased. after some time the direction of co2 diffusion changed, causing the oil of the separator to swell, a condition that continued after the end of the flooding. figure 4 presents oil production curves for flooding with co2-enriched water with and without correction for swel l ing. compared to the water flooding, the amount of oil produced during flooding with co2-enriched water was small. total oil production was 0.047 times the pore volume equivalent to 6.1% of the oil originally present in the sample, if swelling correction is applied, and 0.032 times the pore volume or 4.2% of the oil originally present in the sample, if swelling correction is omitted. the swelling correction is considered valid, and hence the corrected values are preferred. 19 0.1 0.2 0.4 0.3 0.5 0.6 0.7 0.8 10 20 30 40 50 60 700 pr od uc ed o il (m l) fl ow r at e (m l/h ) 0 time (days) 4 5 2 3 1 –1 –2 –3 0 produced oil before swelling correction swelling correction curve produced oil after swelling correction flow rate –0.01 0.01 0.02 0.03 0.04 0.05 0 pr od uc ed o il (f ra ct io n of p or e vo lu m e) –0.03 –0.02 0 1 2 3 4 6 7 8 9 105 0 0.01 pe rm ea bi lit y to w at er ( m ill id ar cy ) 0.02 0.03 0.04 0.05 0.06 injected water volume (fraction of pore volume) –0.02 –0.01 produced oil before swelling correction produced oil after swelling correction model of produced oil permeability from inlet rate fig. 4. produced oil volume (np) versus injected fluid volume (vinj) during flooding with co2-enriched water compared to production curves with and without correction for swelling. model before breakthrough: np = 0. model after breakthrough: np = 0.05 – 0.08 × exp(–vinjj/3.8). fluid saturation (fraction of pore volume): water saturation before co2 flood = 0.719. with swelling correction: water saturation after co2 flood = 0.767, oil production during co2 flood = 0.047. without swelling correction: water saturation after co2 flood = 0.752, oil production during co2 flood = 0.032. fig. 3. swelling correction curve for flooding with co2-enriched water with production curves before and after correction. rosa_2008:rosa-2008 01/07/09 15:47 side 19 a model has been developed that approximates the oil production profile (fig. 4). the fit is reasonable, although inferior to the fit obtained for the water flooding (fig. 2). the inferior fit is at least partly caused by the smaller volume changes compared to the water flooding, causing a greater relative experimental uncertainty. dissolution effects the bulk volume of the sample was reduced by 1.64 ml during the experiment (table 1), equivalent to a grain volume loss of 1.17 ml. this is attributed to a distinct dissolution structure that was found at the inlet end face after the experiment (fig. 5). the structure was due to dissolution of the chalk by the co2-enriched water. dissolution structures were not visible on the other surfaces of the sample. porosity of the sample increased by 0.32% of the bulk volume during the experiment (table 1), which indicates that 1.01 ml of grain material were dissolved from the interior of the sample during the experiment. combining the bulk volume loss and the porosity increase indicates a total dissolution of 2.18 ml of grain material, which is equivalent to the loss of 5.91 g of calcite. the measured weight loss was 4.32 g, which is considered to agree within the experimental uncertainty (table 1). the gas permeability increased by 23% during the experiment (table 1), which is remarkable compared to the modest increase in porosity. the increase in both porosity and permeability was evenly distributed among the four subsamples, indicating a certain amount of dissolution throughout the sample (olsen 2007). the evidence shows that dissolution took place both close to where the co2-enriched water entered the sample and within the sample. a small length reduction (table 1) occurred due to the formation of the dissolution structure. conclusions water flooding of a low-permeable ekofisk chalk sample from the nana-1x well in the danish north sea resulted in an oil saturation of 28.0% of the pore volume with an oil recovery of 63.8% of the oil originally present in the sample. flooding with co2-enriched water corresponding to 10.4 times the pore volume increased the oil recovery by 6.1% of the oil originally present in the sample, resulting in a final oil saturation of 23.3% of the pore volume. compared to the water flooding, the flooding with co2-enriched water increased the final oil recovery by 9.6%. a correction procedure was applied to correct for diffusion processes within the separator of the rig. at the inlet end of the sample a dissolution structure was created by the co2-enriched water. the rest of the sample was visually unaffected by the flooding, but the injection resulted in a 1.1% increase in porosity and a 23% increase in gas permeability, which shows that some dissolution took place within the pore space of the sample. acknowledgements financial support was received from the ministry of science, technology and innovation, and from the european network of excellence on geological storage of co2 that is co-funded by the european com mission within the 6th framework programme. references chang, y., coats, b.k. & nolen, j.s. 1998: a compositional model for co2 floods including co2 solubility in water. spe reservoir evaluation & engineering 1, 155–160. jørgensen, l.n. 1992: dan field – denmark, central graben, danish north sea. in: foster, n.h. & beaumont, e.a. (eds): atlas of oil and gas fields. structural traps vi, 199–218. tulsa: american association of petroleum geologists. olsen, d. 2007: increased oil recovery from the danish north sea chalk fields. flooding experiment ocd1. danmarks og grønlands geolo giske undersøgelse rapport 2007/30, 20 pp. . 20 author’s address geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: do@geus.dk 1 cm fig. 5. the inlet end of the composite sample showed a dissolution structure after flooding with co2-enriched water. the star-like shape is governed by the arrangement of the distributor channels in the end pieces of the core holder. rosa_2008:rosa-2008 01/07/09 15:47 side 20 geological survey of denmark and greenland bulletin 13, 2007, 01-07 geological survey of denmark and greenland bulletin 13 • 2007 review of survey activities 2006 edited by martin sønderholm and a.k. higgins geological survey of denmark and greenland danish ministry of the environment geological survey of denmark and greenland bulletin 13 keywords geological survey of denmark and greenland, survey organisations, current research, denmark, greenland. cover photographs from left to right 1. study of potentially gold-bearing archaean supracrustal rocks south-west of the isua supracrustal belt, southern west greenland. photo: adam a. garde. 2. detailed surveying of the coastal zone is an important task during construction of e.g. harbours, bridges and man-made beach resorts. photo: merete binderup. 3. drilling of the enreca-2 well on the vietnamese island of phu quoc. the well encountered a 500 m thick lower cretaceous fluvial sandstone unit that is widely distributed in the area including onshore cambodia. the formation constitutes a potential reser voir for hydrocarbon and freshwater in the area. photo: lars henrik nielsen. 4. mapping of hydraulic fractures induced for enhanced treatment of contaminated soil at the kluzcewo airport in poland. the acti vities were carried out in connection with the fp6 project stresoil sponsored by the eu. photo: tomasz kasela. frontispiece: facing page study of oil seeps in basalts on the south-western shores of nuussuaq, west greenland during an excursion in 2004. petroleum exploration offshore west greenland is now going into a new phase after the licensing rounds in 2006 and 2007 where seven new licences were awarded. photo: martin sønderholm. chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, department of earth sciences, university of aarhus; minik rosing, geological museum, university of copenhagen; finn surlyk, department of geography and geology, university of copenhagen scientific editors: martin sønderholm and a.k. higgins editorial secretaries: jane holst and esben w. glendal referees: (numbers refer to first page of reviewed article): anonymous (29), morten bjerager (9), lars nielsen (17), jens konnerupmadsen (37, 41), kristine thrane (45, 49, 53), michael schultz rasmussen (61), peter engesgaard (65), department of geography and geology, university of copenhagen, denmark. asger k. pedersen (37, 41, 49, 53), geological museum, university of copenhagen, denmark. karen luise knudsen (21), department of earth sciences, university of aarhus, denmark. claus heinberg (21), roskilde university, denmark. henrik tirsgaard (9), poul henrik larsen (73), steve dorobek (73), maersk oil and gas, copenhagen, denmark. niels l. westphal (13), michael larsen (17), gregers dam (25, 29, 33), dong energy, hørsholm, denmark. graham pearson (45), durham university, uk. jens havskov (57), university of bergen, norway. illustrations: stefan sølberg with contributions from jette halskov and christian rasmussen lay-out and graphic production: annabeth andersen printers: schultz grafisk, albertslund, denmark manuscripts submitted: 2–19 february 2007 final versions approved: 14 august 2007 printed: 12 october 2007 issn 1603-9769 (review of survey activities) issn 1604-8156 (geological survey of denmark and greenland bulletin) isbn 978-87-7871-202-8 available from geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk or geografforlaget aps, filosofgangen 24,1., dk-5000 odense c, denmark phone: +45 63 44 16 83, fax: +45 63 44 16 97, e-mail: go@geografforlaget.dk © de nationale geologiske undersøgelser for danmark og grønland (geus), 2007 for the full text of the geus copyright clause, please refer to www.geus.dk/publications/bull 4 7. review of survey activities 2006 k. sørensen 9. chalk depth structure maps, central to eastern north sea, denmark o.v. vejbæk, t. bidstrup, p. britze, m. erlström, e.s. rasmussen and u. sivhed 13. are carboniferous coals from the danish north sea oil-prone? h.i. petersen and h.p. nytoft 17. prediction of reservoir sand in miocene deltaic deposits in denmark based on high-resolution seismic data e.s. rasmussen, t. vangkilde-pedersen and p. scharling 21. environmental change in danish marine waters during the roman warm period inferred from mollusc data p. rasmussen, k.s. petersen and d.b. ryves 25. petroleum systems and structures offshore central west greenland: implications for hydrocarbon prospectivity u. gregersen, t. bidstrup, j.a. bojesen-koefoed, f.g. christiansen, f. dalhoff and m. sønderholm 29. provenance of cretaceous and paleocene sandstones in the west greenland basins based on detrital zircon dating a. scherstén and m. sønderholm geus working areas 2006. orange areas are covered in this volume. for further information on other working areas please refer to our website: www.geus.dk/international 5 33. a multi-disciplinary study of phanerozoic landscape development in west greenland j.m. bonow, p. japsen, p.f. green, r.w. wilson, j.a. chalmers, k.e.s. klint, j.a.m. van gool, k. lidmar-bergström and a.k. pedersen 37. pre-metamorphic hydrothermal alteration with gold in a mid-archaean island arc, godthåbsfjord, west greenland a.a. garde, h. stendal and b.m. stensgaard 41. gold-hosting supracrustal rocks on storø, southern west greenland: lithologies and geological environment c. knudsen, j.a.m. van gool, c. østergaard, j.a. hollis, m. rink-jørgensen, m. persson and k. szilas 45. p–t history of kimberlite-hosted garnet lherzolites from south-west greenland m.t. hutchison, l.j. nielsen and s. bernstein 49. two tectonically significant enclaves in the nordre strømfjord shear zone at ataneq, central west greenland w.e. glassley, j.a. korstgård and k. sørensen 53. a well-preserved bimodal archaean volcanic succession in the tasiusarsuaq terrane, south-west greenland h. stendal and a. scherstén 57. seismic hazard assessment of greenland p. voss, s.k. poulsen, s.b. simonsen and s. gregersen 61. development of marine landscape maps for the baltic sea and the kattegat using geophysical and hydrographical parameters z.k. al-hamdani, j. reker, j.o. leth, a. reijonen, a.t. kotilainen and g.e. dinesen 65. shallow groundwater quality in latvia and denmark e. gosk, i. levins and l.f. jørgensen 69. bayesian belief networks as a tool for participatory integrated assessment and adaptive groundwater management: the upper guadiana basin, spain h.j. henriksen, p. rasmussen, j. bromley, a. de la hera portillo and m.r. llamas 73. cenozoic evolution of the vietnamese coastal margin m.b.w. fyhn, l.h. nielsen and l.o. boldreel 7 the geological survey of denmark and greenland (geus) has lived through a period of intense unrest during most of 2006. this was a consequence of the decision by the danish government in 2005 to carry out a major reorganisation of the danish research world. the aim was to improve the quality and competitiveness of danish universities and research organisations (such as geus), by means of a fusion of universities and the merging of independent research institutions with the universities. at the conclusion of this process all but three of the major government research institutions were merged with the universities in århus and copenhagen, and the technical university in lyngby. for many reasons, one being the special tasks that geus is responsible for in greenland, the government decided that geus should continue as one of these three independent national research institutions. this issue of review of survey activities (rosa) is the fourth published after it was decided to publish an annual research overview illustrating the activities of geus in denmark, greenland and other countries building on a longstanding tradition of a greenland review publication. although only approximately a third of our institution’s turnover is related to greenland, no less than 34 papers out of the 73 published in the four issues of rosa (including this one) relate to greenland activities. greenland thus continues to be a potent measure of scientific productivity at geus, and it is deeply satisfying to acknowledge the dedication to the geology of greenland as one of the major factors that helped secure the independence of the survey. although this issue of review of survey activities presents 17 papers providing a panorama of the current research carried out at geus, it illustrates only a small part of the wide range of projects undertaken in denmark and greenland and other countries in 2006. a factual overview of the activities of geus as a whole can be seen on the geus website. in the present volume, three papers deal with various aspects of petroleum exploration in the danish part of the north sea. one of these papers links exploration for deep groundwater aquifers onshore denmark with offshore petroleum exploration. another paper from denmark is devoted to recent environmental changes during the roman warm period. in greenland, exploration activities for both minerals and petroleum have reached unprecedented highs during 2006. this is reflected by the nine papers dealing with projects related to greenland. three papers are concerned with petroleum geological matters; two of these are a result of seismic and provenance studies carried out by geus for the government of greenland in preparation for the disko west 2006 licensing round. the third paper presents a combination of landscape and fission track analysis to elucidate the cenozoic uplift of the west greenland margin. five papers are related to various aspects of mineral exploration, focussing on gold (two papers), diamonds (one paper) and general mapping of proterozoic and archaean primary geological environments with special emphasis on tectonic and mineralising events (two papers). the last article on greenland deals with seismic hazard assessment, a discipline carried out at geus since the incorporation of the seismological service for denmark and greenland in 2004. geus is also involved in a wide range of activities outside its core working areas in denmark and greenland that are well illustrated by the four last papers in this review. three papers are related to national implementation of european legislation, such as the water framework directive. the first paper documents a circum-baltic project on the development of marine landscape maps, and the two following papers deal with various aspects of groundwater management. one of these is related to the establishment of a groundwater monitoring programme in latvia, while the other deals with stakeholder participation in adaptive and integrated water resource management in spain. the last paper on the cenozoic evolution of the vietnamese offshore region illustrates how the survey’s broad technical expertise has been put to use in developing countries. review of survey activities 2006 kai sørensen director © geus, 2007. geological survey of denmark and greenland bulletin 13, 7 only. available at: www.geus.dk/publications/bull geological survey of denmark and greenland bulletin 13, 2007, 25-28 a detailed geophysical mapping project has been carried out by the geological survey of denmark and greenland (geus) in the offshore region south-west and west of disko and nuussuaq, central west greenland as part of the preparations for the disko west licensing round in 2006 (fig. 1). the main purpose of the study was to evaluate the prospectivity of this almost 100 000 km2 large region, and to increase knowledge of basin evolution and the structural development. results of the work, including a new structural elements map of the region and highlights of particular interest for hydrocarbon exploration of this area, are summarised below. evidence of live petroleum systems has been recognised in the onshore areas since the beginning of the 1990s when seeps of five different oil types were demonstrated (bojesenkoefoed et al. 1999). oil seeps suggesting widely distributed marine source rocks of mesozoic age are particularly promising for the exploration potential (bojesen-koefoed et al. 2004, 2007). furthermore, possible dhis (direct hydro carbon indicators) such as gas-clouds, pock marks, bright spots and flat events have been interpreted in the offshore region (skaarup et al. 2000; gregersen & bidstrup in press). the evaluation of the region (fig. 1) is based on all public and proprietary seismic data together with public domain mag netic and gravity data. the seismic data (a total of c. 28 000 line km) are tied to the two existing offshore exploration wells in the region (hellefisk-1 and ikermiut-1). the study also incorporates information on sediments and volcanic rocks from onshore disko and nuussuaq (fig. 2). ten seismic horizons ranging from ‘mid-cretaceous’ to ‘base quaternary’ (fig. 2) have been interpreted regionally. large correlation distances to wells, varying data quality and a thick cover of basalt in the north-eastern part of the region, add uncertainty in the regional interpretation, especially for the deeper horizons such as the ‘mid-cretaceous’ equivalent to santonian sandstone interval drilled in qulleq-1 far south. based on the seismic interpretation (fig. 3) structural elements maps, horizon-depth maps and isopach maps have been produced; these maps, together with general stratigraphic knowledge on potential reservoirs, seals and source rocks (fig. 2), provide important information for discussions of critical play elements including kitchens and structures. the existence of many large structures combined with the evidence of live petroleum systems has spurred the recent major interest for hydrocarbon exploration in the region. structural development and basin evolution a number of deep basins with cretaceous and cenozoic sedimentary successions have been recognised offshore west greenland since the 1970s (e.g. chalmers et al. 2001). in early to mid-cretaceous times a number of major structural complexes and basins developed in the region (figs 1, 3), mainly as a result of extensional faulting. these include the © geus, 2007. geological survey of denmark and greenland bulletin 13, 25–28. available at: www.geus.dk/publications/bull 25 petroleum systems and structures offshore central west greenland: implications for hydrocarbon prospectivity ulrik gregersen, torben bidstrup, jørgen a. bojesen-koefoed, flemming g. christiansen, finn dalhoff and martin sønderholm fig. 1. structural elements offshore the disko–nuussuaq region, central west greenland. the position of the seismic example in fig. 3 is shown. aasiaat structural trend, the kangerluk structure, the aasiaat basin, the sisimiut basin and the nagssugtôq subbasin (fig. 1). deep-seated fault-bounded basins lo cally showing anticlinal structures are occasionally observed below the interpreted ‘mid-cretaceous’ seismic horizon. seabed sampling has shown the presence of ordovician carbonate strata on the davis strait high (dalhoff et al. 2006). together with reworked jurassic or older palyno morphs observed in the qulleq-1 well farther south (nøhrhansen et al. 2000) these suggest the possibility of pre-cretaceous strata in the deepest parts of these basins. based on outcrop studies, regional cretaceous sand-prone units are expected to be present in the offshore region, both as deltaic and shallow-marine deposits and as turbidite deposits (fig. 2). during the late cretaceous a new rifting episode was initiated and was characterised by normal faulting, subsidence and syn-rift sedimentation including deposition of a thick mudstone succession of campanian–maa strich tian age (dam et al. 2000). this phase lasted into early paleocene times resulting in repeated erosion and filling of subaerial valley and submarine canyons and formation 26 fig. 2. simplified stratigraphic scheme with lithology from present coastal areas of disko and nuussuaq (north), and from offshore west greenland wells towards the south. the main phases of tectonism and subsidence are listed. the main interpreted seismic horizons are seabed, base quaternary (bq), midmiocene (mm), lower eocene (le), top basalt 3 (tb3), upper paleocene (up), top basalt 2 (tb2), top basalt 1/base basalt 2 (tb1), top cretaceous (tc), mid-cretaceous (mc) and unspecified deep reflections. the most likely intervals with source rocks (so), reservoir sands (r) and seals (se) are also indicated. fig. 3. seismic section (ggui95-17) through eastern parts of the study region, showing main structural elements. the main structural elements are from west to east: the easternmost parts of the aasiaat basin, the ilulissat graben edge (ige), the ilulissat graben on both sides of the ilulissat high, and the westernmost part of the disko high. note the amplitude anomalies above the ilulissat high (see fig. 1 for location). of regional unconformities (dam & sønderholm 1998; dam 2002). a zircon age provenance study of sandstone units in outcrops and offshore exploration wells indicates that most of the sand units show local age sig natures characteristic of the green land shield to the east. however, a gren ville age component of probable can a dian derivation also seems to be present in the deep-water deposits of the nuus suaq basin and in the qulleq-1 well suggesting a long-shore transport component (scherstén & sønderholm 2007 – this volume). during the paleocene–eocene a major episode of volcanic eruption took place, and some hundreds of metres of thick basalts cover part of the offshore region (fig. 3). the basalts may reach thicknesses of more than 2 km in the north-eastern offshore part of the study region. however, the current mapping suggests that the volcanic succession is thinner and less widely distributed than suggested in previous publications and maps (e.g. chalmers et al. 1993; chalmers & pulvertaft 2001; skaarup 2002). strike-slip movements during the late paleocene and early eocene caused local transpression of structures primarily along the ikermiut fault zone, and locally in the basins and structures farther north, contemporaneous with subsidence in the ikermiut basin region. transtensional and extensional movements farther to the north-east, subsequent to the extrusion of the paleocene basalts, resulted in the development of the more than 200 km long ilulissat graben (fig. 1). during the eocene, and especially during the late mio cene to pliocene, the offshore basins subsided rapidly, and large sedimentary wedges prograded towards the west and south, possibly as a consequence of neogene uplift in the pres ent onshore areas to the east (fig. 3; dalhoff et al. 2003; japsen et al. 2005; bonow et al. 2007 – this volume). petroleum systems and prospectivity based on seismic interpretation, depth conversion using seismic velocities, sonic log data from the wells and the thermal maturity gradient from selected wells, the most likely source rock intervals (mid-cretaceous and lower paleocene) seem to be mature in large parts of the region (fig. 4), though seismic interpretation is difficult. in particular the aasiaat basin, the aasiaat structural trend, the north ungava basin, the ikermiut basin, the sisimiut basin and the ilulissat graben (fig. 1) may have adequate dimensions and depths to have potential as kitchens for hydrocarbon generation, with the potential also depending on factors such as source rocks being present in sufficient quality and quantity. this study indicates that the interpreted 27 fig. 4. simplified prospectivity map. cretace ous to palaeogene structures and major midcretaceous 4-way dip closures to the west, pos sible hydrocarbon migration pathways and mid-cretaceous hydrocarbon generation areas (maturity levels – early to main oil: ~0.5 – >1% ro; late oil: ~1 – >1.3% ro; gas: >1.3% ro). 28 source rock intervals possibly came into the oil window after mid-miocene time, subsequent to the formation of the main structural closures providing a favourable timing for charging. in the interpreted cretaceous and cenozoic sections, amplitude anomalies are locally observed and may be interpreted as dhis (such as e.g. bright spots above the ilulissat high in fig. 3) that could be caused by trapped hydrocarbon. clusters of dhis are located especially over or near the supposed cretaceous kitchen areas, and also locally where satellite slicks have been recorded (fig. 4), and contribute to an indication of live petroleum systems in the offshore region. mapping of the cretaceous and palaeogene intervals and structural highs has revealed many structures. large structural closures can be outlined both in the western part of the region (in the aasiaat basin, the aasiaat structural trend and the kangerluk structure), in the eastern part of the region (both along the edge and within the ilulissat graben) and in the southern part of the region related to the ikermiut fault zone (fig. 4). the cretaceous and palaeogene structural closures are situated close to supposed kitchen areas (fig. 4), and together with the oil seeps and reservoir quality sandstones known onshore and their supposed offshore equivalents, these elements indicate that the offshore area west and south of disko could potentially be prospective. acknowledgements the geophysical study was supported by the bureau of minerals and petroleum, government of greenland. tgs-nopec and nunaoil a/s are thanked for permission to publish the structural maps that incorporate proprietary seismic and satellite slick data. references bojesen-koefoed, j.a., christiansen, f.g., nytoft, h.p. & pedersen, a.k. 1999: oil seepage onshore west greenland: evidence of multiple source rocks and oil mixing. in: fleet, a.j. & boldy, s.a.r. (eds): petroleum geology of northwest europe: proceedings of the 5th conference, 305–314. london: geological society. bojesen-koefoed, j.a., nytoft, h.p. & christiansen, f.g. 2004: age of oils in west greenland: was there a mesozoic seaway between greenland and canada? geological survey of denmark and greenland bulletin 4, 49–52. bojesen-koefoed, j.a., bidstrup, t., christiansen, f.g., dalhoff, f., greger sen, u., nytoft, h.p., nøhr-hansen, h., pedersen, a.k. & sønder holm, m. 2007: petroleum seepages at asuk, disko, west greenland – implications for regional petroleum exploration. journal of petroleum geo logy 30, 219–236. bonow, j.m., japsen, p., green, p.f., wilson, r.f., chalmers, j.a., klint, k.e., van gool, j.a.m., lidmar-bergström, k. & pedersen, a.k. 2007: a multi-disciplinary study of phanerozoic landscape development in west greenland. geological survey of denmark and greenland bulletin 13, 41–44. chalmers, j.a. & pulvertaft, t.c.r. 2001: development of the continental margins of the labrador sea – a review. in: wilson, r.c.l. et al. (eds): non-volcanic rifting of continental margins: a comparison of evidence from land and sea. geological society special publication (london) 187, 79–107. chalmers, j.a., pulvertaft, t.c.r., christiansen, f.g., larsen, h.c., laursen, k.h. & ottesen, t.g. 1993: the southern west greenland continental margin: rifting history, basin development, and petroleum potential. in: parker, j.r. (ed.): petroleum geology of nw europe: proceedings of the 4th conference, 915–931. london: geological society. chalmers, j.a., christiansen, f.g., sønderholm, m., olsen, j.c., mykle bust, r. & schønwandt, h.k. 2001: geological information base growing on north atlantic rift basins. data developed for greenland licensing. offshore 61(11), 87–89, 100. dalhoff, f., chalmers, j.a., gregersen, u., nøhr-hansen, h., rasmussen, j.a. & sheldon, e. 2003: mapping and facies analysis of paleo cene–mid-eocene seismic sequences, offshore southern west green land. marine and petroleum geology 20, 935–986. dalhoff, f., larsen, l.m., ineson, j.r., stouge, s., bojesen-koefoed, j.a., lassen, s., kuijpers, a., rasmussen, j.a. & nøhr-hansen, h. 2006: continental crust in the davis strait: new evidence from seabed sampling. geological survey of denmark and greenland bulletin 10, 33–36. dam, g. 2002: sedimentology of magmatically and structurally controlled outburst valleys along rifted volcanic margins; examples from the nuussuaq basin, west greenland. sedimentology 49, 505–532. dam, g. & sønderholm, m. 1998: sedimentological evolution of a faultcontrolled early paleocene incised-valley system, nuussuaq basin, west greenland. in: shanley, k.w. & mccabe, p.j. (eds): relative role of eustasy, climate, and tectonism in continental rocks. society for sedimentary geology (sepm) special publication 59, 109–121. dam, g., nøhr-hansen, h., pedersen, g.k. & sønderholm, m. 2000: sedimentary and structural evidence of a new early campanian rift phase in the nuussuaq basin, west greenland. cretaceous research 21, 127–154. gregersen, u. & bidstrup, t. in press: structures and hydrocarbon prospectivity in the northern davis strait area, offshore west greenland. petroleum geoscience. japsen, p., green, p.f. & chalmers, j.a. 2005: separation of palaeogene and neogene uplift on nuussuaq, west greenland. journal of the geological society (london) 162, 299–314. nøhr-hansen, h., piasecki, s., rasmussen, j.a. & sheldon, e. 2000: biostratigraphy of well 6354/4-1 (qulleq-1), west greenland. dan marks og grønlands geologiske undersøgelse rapport 2000/101, 81 pp. scherstén, a. & sønderholm, m. 2007: provenance of cretaceous and paleocene sandstones in the west greenland basins based on detrital zircon dating. geological survey of denmark and greenland bulletin 13, 37–41. skaarup, n. 2002: evidence for continental crust in the offshore palaeo gene volcanic province, central west greenland. geology of green land survey bulletin 191, 97–102. skaarup, n., chalmers, j.a. & white, d. 2000: an avo study of a possible new hydrocarbon play, offshore central west greenland. american association of petroleum geologists bulletin 84, 174–182. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ug@geus.dk geological survey of denmark and greenland bulletin 15, 2008, 25-28 denmark is a key region for studies of the cenozoic development of scandinavia because paleocene to upper miocene sediments crop out across the country and because it is possible to correlate these occurrences with the up to 3 km thick cenozoic succession of the north sea basin. however, the reason why the cenozoic deposits occur close to the surface of the earth in denmark is that the sediments have been exhumed from their cover of younger rocks. this implies that a reconstruction of the cenozoic development across denmark – involving both burial and exhumation – must rely on sedimentological and seismic studies of preserved sediments as well as on physical parameters that may yield evidence of the postdepositional history of the sediments now at the surface. only if the burial and exhumation history of the basins can be deciphered is it possible to infer the geological development in the scandinavian hinterland where cenozoic sediments are rarely preserved. we have identified four mesozoic–cenozoic palaeothermal phases related to burial and subsequent exhumation, and one phase reflecting climate change during the eocene. this is based on new apatite fission-track analyses (afta) and vitrinite reflectance data from eight danish wells (japsen et al. 2007a). the study combined thermal history reconstruction with exhumation studies based on palaeoburial (sonic velo city), stratigraphic and seismic data (cf. japsen & bidstrup 1999; green et al. 2002; nielsen 2003; rasmussen 2004; japsen et al. 2007b). two of the exhumation phases occurred during the mid-jurassic and the mid-cretaceous. in this study we focus on the cenozoic development and on the early and late neogene exhumation phases during which up to 1 km of sediments were removed across most of the danish region (fig. 1). early and late neogene exhumation a major phase of exhumation of the parts of the eastern north sea basin adjacent to the presently exposed basement areas in norway and sweden began between 30 and 20 ma ago according to the afta data (fig. 1). we suggest that this phase corresponds to the oligocene–miocene unconformity (c. 24 ma). prior to this early neogene exhumation phase, the mesozoic sediments in the felicia-1 well were at maximum burial. the maximum burial of the mesozoic sediments in the hans-1 well occurred during the mid-cretaceous, prior to inversion along the sorgenfrei–tornquist zone, whereas maximum burial during the cenozoic occurred prior to the early neogene exhumation phase. late neogene exhumation began between 10 and 5 ma ago according to the afta data from e.g. the års-1 and felicia-1 wells. only one sample limits the onset of this phase to begin before 5 ma, and the onset of the exhumation probably began in the early pliocene at c. 4 ma as suggested by the prominent unconformity of this age. after 4 ma significant progradation from scandinavia into the north sea basin and increased subsidence in the central north sea were initiated. a section of 450–850 m was removed in the central part of the norwegian–danish basin in association with this reshaping of the north sea basin. the exhumation affected extensive regions where neogene strata are truncated, and in southern norway and sweden a sub-cretaceous etched surface was reexposed along the coasts (lidmar-bergström et al. 2000). cenozoic palaeogeography and isochores predating the neogene exhumation of the eastern north sea basin peter japsen, erik s. rasmussen, paul f. green, lars henrik nielsen and torben bidstrup © geus, 2008. geological survey of denmark and greenland bulletin 15, 25–28. available at: www.geus.dk/publications/bull 25 fig. 1. thickness of the section removed during cenozoic exhumation. maximum burial of the mesozoic sediments took place during the late neogene in most of the study area. near the present coasts of norway and sweden, however, maximum burial occurred subsequent to deposition of the oligocene wedges and prior to early neogene exhumation. only in the hans-1 well did maximum burial occur during the mesozoic. modified from japsen et al. (2007a). 6°e 10°e 12°e58°n 57°n 56°n 750 500 250 contour interval 250 m well data well, afta data cenozoic exhumation 100 km late neogene early neogene maximum burial felicia-1 års-1 hans-1 maps of palaeogeography and isochores prior to the neogene exhumation phases we have compiled a series of maps to illustrate the cenozoic development of the eastern north sea basin (figs 2–6). the isochore maps are based on present thicknesses and estimates of the removed sections constrained by the total amount of section removed by cenozoic exhumation. the estimates of the removed sections of specific ages are furthermore constrained by the known geology and our interpretation of the depositional pattern. each isochore map shows the thicknesses prior to the first phase of exhumation that affected the distribution of the unit; i.e. prior to early neogene exhumation (starting at c. 24 ma) in the case of the palaeogene units and prior to late neogene exhumation (starting at c. 4 ma) for the neogene units. the palaeogeographical maps show the distribution of onshore and offshore areas separated by the coastline based on the known geology. in the areas where the deposits of a given age have been removed, the map has been drawn from an estimate based on our interpretation of the depositional system and on the inferred burial and exhumation history. in the early eocene, the study region was covered by a deep sea as demonstrated by hemipelagic deep-marine sedimentation that lasted until latest eocene times (fig. 2a; heilmann-clausen et al. 1985; michelsen et al. 1998). parts of scandinavia were probably covered by eocene sediments because eocene deposits within the norwegian–danish ba sin contain a deep marine fauna close to the sorgen frei–tornquist zone with no signs of a near-shore fauna (c. heilmann-clausen, personal communication 2005). further more, reworked eocene dinocysts and clasts of eocene muds interbedded with miocene deltaic sediments in denmark clearly indicate the presence of marine eocene deposits in the scandinavian hinterland (rasmussen 2004). the thickness of the upper paleocene – eocene sediments that was removed during the neogene rise of scandinavia was probably limited (<200 m; fig. 2b). during the oligocene, major clastic wedges prograded into the norwegian–danish basin from present-day norway (fig. 3a; michelsen et al. 1998; clausen et al. 1999; faleide et al. 2002). this change in depositional environment from eocene times indicates a phase of tectonic uplift of southern norway (c. 33 ma) (e.g. michelsen et al. 1998; lidmar-berg ström et al. 2000; faleide et al. 2002). this tectonic activity is shown by reactivation of faults on the ringkøbing–fyn high and by movements of salt structures in the nor we gian–danish basin (rasmussen 2004). we suggest that a thick succession of oligocene shelf and delta sediments that 26 b 100 km a b a present coastline sediment transport 200–300 100–200 <100 >1000 800–1000 500–800 300–500 thickness (m)succession truncated, estimated thickness succession absent, estimated thickness onshore offshore 50 ma 25 ma fig. 2. a: early eocene palaeogeography, 50 ma. b: upper paleocene – eocene isochore prior to neogene exhumation. modified from japsen et al. (2007a). fig. 3. a: late oligocene palaeogeography, 25 ma. b: oligocene isochore prior to neogene exhu mation. modified from japsen et al. (2007a). legend to figs 2–6. prograded southwards from southern norway and westwards from southern sweden made up a substantial part of the section that was removed during the neogene exhumation at the locations of the felicia-1 and hans-1 wells, respectively (>1 km) (fig. 3b). during the early miocene, deltas prograded to the south and south-west and reached a thickness up to 300 m (fig. 4a). the deposition of coarse-grained sediments reached the southern part of the norwegian–danish basin and the ring købing–fyn high during the neogene (rasmussen 2004). these prograding systems reflect a redistribution of sediments caused by uplift in the scandinavian hinterland that resulted in the early neogene exhumation near the craton. the occurrence of both immature and mature sediments reflects erosion of weathered as well as newly exposed basement. in the danish area, the early miocene tectonic activity is demonstrated by coarse-grained, braided fluvial systems and later by a sudden increase in the heavy mineral content at c. 17 ma (rasmussen 2004). the thickness of the lower – lower middle miocene sediments that was removed during late neogene exhumation was probably limited (<500 m), because mainly the onshore part of these deposits were af fected by exhumation (fig. 4b). in the middle miocene, at c. 15 ma, a distinct marine flooding of the area took place, and up to 150 m of clayey sediments were deposited in south-western denmark during the middle–late miocene (fig. 5a) (rasmussen 2005). the flooding was caused partly by eustatic sea-level rise related to the mid-miocene climatic optimum and partly by increased subsidence in the eastern north sea basin in the late miocene (rasmussen 2004). the thickness of the upper middle – lower upper miocene sediments that was removed during late neogene exhumation was limited (<200 m; fig. 5b). at the end of the miocene resumed delta progradation from north-east and in particular the east occurred (fig. 6a) (rasmussen 2005). the infilling of the north sea basin continued during the pliocene where up to 500 m were depos ited in the central north sea basin. substantial thicknesses (<500 m) of upper upper miocene – pliocene sediments were removed during late neogene exhumation across an extensive region of the eastern north sea (fig. 6b). summary and implications this study emphasises that the tectonic development of a region cannot be reconstructed solely on evidence from the sedimentary record. such reconstructions need also to be based on the record of physical indicators of palaeothermal and palaeoburial phases related to the former presence of geological units now removed. we have found evidence for three tectonic phases that have affected southern scandinavia during the cenozoic: (1) a phase that began at the eocene–oligocene transition at c. 33 ma as indicated by the onset of progradation of clastic wedges away from southern norway and 27 b a b a ? ? 15 ma 12 ma fig. 4. a: middle miocene palaeogeography, 15 ma. b: lower – lower middle miocene isochore prior to late neogene exhumation. modified from japsen et al. (2007a). fig. 5. a: late miocene palaeogeography, 12 ma. b: upper middle – lower upper miocene isochore prior to late neogene exhumation. modified from japsen et al. (2007a). inferred progradation away from southern sweden. (2) a phase that began at the oligocene–miocene transition at c. 24 ma, as indicated by early neogene exhumation of the areas adjacent to the presently exposed basement areas and by early miocene coarse-grained braided fluvial systems south of scandinavia. (3) a phase that began in the early pliocene at c. 4 ma, as indicated by the widespread, late neogene exhumation, the intra-pliocene unconformity and subsequent tilting of the neogene succession in the eastern north sea. these phases are consistent with the stratigraphy in the north-east atlantic ocean (stoker et al. 2005) and around southern norway (e.g. michelsen et al. 1998; faleide et al. 2002; rasmussen 2004). the observations thus suggest that southern norway, with peaks higher than 2 km above sea level, emerged during several phases since eocene times, when a deep ocean covered much of the region. references clausen, o.r., gregersen, u., michelsen, o. & sørensen, j.c. 1999: factors controlling the cenozoic sequence development in the eastern parts of the north sea. journal of the geological society of london 156, 809–816. faleide, j.i., kyrkejbø, r., kjennerud, t., gabrielsen, r.h., jordt, h., fanavoll, s. & bjerke, m. 2002: tectonic impact on sedimentary processes during cenozoic evolution of the northern north sea and surrounding areas. in: doré, a.g. et al. (eds): exhumation of the north atlantic margin: timing, mechanisms and implications for petroleum exploration. geological society special publications (london) 196, 235–269. green, p.f., duddy, i.r. & hegarty, k.a. 2002: quantifying exhumation from apatite fission-track analysis and vitrinite reflectance data: precision, accuracy and latest results from the atlantic margin of nw europe. in: doré, a.g. et al. (eds.): exhumation of the north atlantic margin: timing, mechanisms and implications for petroleum exploration. geo logical society special publications (london) 196, 331–354. heilmann-clausen, c., nielsen, o.b. & gersner, f. 1985: lithostratigraphy and depositional environments in the upper palaeocene and eocene of denmark. bulletin of the geological society of denmark 33, 287–323. japsen, p. & bidstrup, t. 1999: quantification of late cenozoic erosion in denmark based on sonic data and basin modelling. bulletin of the geological society of denmark 46, 79–99. japsen, p., green, p.f., nielsen, l.h., rasmussen, e.s. & bidstrup, t. 2007a: mesozoic–cenozoic exhumation in the eastern north sea ba sin: a multi-disciplinary study based on palaeo-thermal, palaeo-burial, stratigraphic and seismic data. basin research 19, 451–490. japsen, p., mukerji, t. and mavko, g. 2007b: constraints on velocitydepth trends from rock physics models. geophysical prospecting 55, 135–154. lidmar-bergström, k., ollier, c.d. & sulebak, j.c. 2000: landforms and uplift history of southern norway. global and planetary change 24, 211–231. michelsen, o., thomsen, e., danielsen, m., heilmann-clausen, c., jordt, h. & laursen, g. 1998: cenozoic sequence stratigraphy in the eastern north sea. in: de graciansky, p.-c. et al. (eds): mesozoic and cenozoic sequence stratigraphy of european basins. society of economic pale ontologists and mineralogists special publication 60, 91 –118. nielsen, l.h. 2003: late triassic – jurassic development of the danish basin and fennoscandian border zone, southern scandinavia. in: ineson, j. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 459–526. rasmussen, e.s. 2004: the interplay between true eustatic sea-level changes, tectonics, and climatical changes: what is the dominating factor in sequence formation of the upper oligocene–miocene succession in the eastern north sea basin, denmark? global and plan etary change 41, 15–30. rasmussen, e.s. 2005: the geology of the upper middle – upper mio cene gram formation in the danish area. palaeontos 7, 5–18. stoker, m.s., praeg, d., shannon, p.m., hjelstuen, b.o., laberg, j.s., nielsen, t., van weering, t.c.e., sejrup, h.p. & evans, d. 2005: neo gene evolution of the atlantic continental margin of nw europe lofoten islands to sw ireland: anything but passive. in: doré, a.g. & vining, b.a. (eds): petroleum geology: north-west europe and global per spectives. proceedings of the 6th petroleum geology conference, 1057–1076. london: geological society. zachos, j.c., pagani, m., sloan, l.c., thomas, e. & billups, k. 2001: trends, rythms, and aberrations in global climate 65 ma to present. science 292, 686–693. authors’ addresses p.j., e.s.r., l.h.n. & t.b., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: pj@geus.dk p.f.g., geotrack international, 37 melville road, brunswick west, victoria 3055, australia. 28 a b 2 ma fig. 6. a: pliocene palaeogeography, 2 ma. b: upper upper miocene – pliocene isochore prior to late neogene exhu mation. modified from japsen et al. (2007a). geological survey of denmark and greenland bulletin 35, 2016, 103-106 103© 2016 geus. geological survey of denmark and greenland bulletin 35, 103–106. open access: www.geus.dk/publications/bull during the past 10–15 years, analytical innovations in geochronology have greatly enhanced the application of geochronological data to geological problems. the advances are mainly driven by developments in laser ablation inductively coupled plasma mass spectrometry (la-icpms) which allows for rapid determination of u-th-pb ages of mineral grains in large sample sets. la-icpms has now become the most common tool in the application of zircon geochronology to a host of different geological problems. one of the most regularly used approaches to evaluate complex u-th-pb geochronological data populations is to use a diagram that combines a binned frequency histogram and a probability density distribution plot (pdp), as described by sircombe (2004) and ludwig (2003). this type of diagram is particularly common in sedimentary provenance studies using large sets of age data (morton et al. 1996; pell et al. 1997; rainbird et al. 1997; sircombe 1999; fergusson et al. 2001). it is also useful for the analysis of complex age patterns in metamorphic and igneous rocks. the microsoft excel workbook agedisplay (sircombe 2004) has often been used to produce these diagrams but it is not supported by versions of excel that are newer than 2003. stand-alone software for pc written in java was therefore developed by the geological survey of denmark and greenland (geus). it is based on the same formulae as described in sircombe (2004), thus the name jagedisplay (with j for java), and has additional features and graphic improvements. here we present the setup, operation and capability of the software with examples from single-grain u/pb age data obtained by la-icpms spot analysis. jagedisplay: software for evaluation of data distributions in u-th-pb geochronology tonny b. thomsen, tjerk heijboer and pierpaolo guarnieri open data file connect to database save chart quit jage display x file fig.1. combined 400–2400 ma histogram and pdp chart with calculated maximum probability ages for distinct distribution peaks of zircon populations from a single sample, using coupled 206pb/238u and 207pb/206pb age pairs. a crossover age of 1100 ma is used, where 206pb/238u ages and 207pb/206pb ages are used below and above this age, respectively. a fixed bin size of 50 ma/bin and a concordia filter of ±10% were used. 104104 setup and operation of jagedisplay the jagedisplay package (current version 1.0) for pc includes an executable jar file, a java ms dos batch file, a html readme file, a brief manual and a folder with test data files. the application is available as shareware from the corresponding author, and it is installed by copying the files you receive to any pc with java. jagedisplay supports both comma-separated (csv) and tab-delimited (tab) input text files, with the same table headers and data types as in agedisplay, iolite (paton et al. 2011), and geus’ in-house zirchron software and database. the input file can include age data from one or more samples, which are loaded from the file menu in the main window (fig. 1) using the open data file option. the input file includes four commonly used radiogenic uth-pb age pairs, viz. 206pb/238u, 207pb/235u, 207pb/206pb and 208pb/232th, together with the sample identification, the name or number of individual analyses, the name of the analytical sequence and information about correction for common lead content. output charts can be saved in pdf or png file format through the save chart option in the file menu. the main window consists of a chart area to the left and the selected sample parameters to the right. when an input file has been loaded, the samples available for plotting are found in the file samples field or, if loaded from a database, in the database samples field. to plot one or more of the loaded samples, they are dragged and dropped in their respective selected file items or selected db items fields. the parameters to be used for the chart are selected, and the create new chart button is pressed. combined histograms and probability–density plots for each radiogenic age pair (206pb/238u, 207pb/235u, 207pb/206pb, 208pb/232th) can be produced. in addition, charts combining the 206pb/238u and 207pb/206pb age pairs can be made by manually setting a specific age (denoted cutoff value) that controls the change from using 206pb/238u ages for the younger age range to 207pb/206pb ages for the older age range of the chart (fig. 1). the frequency in terms of number of analyses for each bin in the histogram is shown on the left y-axis, and the calculated probability distribution of zircon ages for the data set on the right y-axis. thus, the probability–density distribution provides an estimate of the true mineral-age distribution for a given rock sample. the plots can be adjusted by the following parameters: • concordia filter is the maximum allowed discordance (in per cent) for analyses to be considered as concordant, and is determined as 100 ± set#value. this separates the age data used for the plot into two different fields representing ‘concordant’ and ‘discordant’ data. in detrital zircon provenance studies a value of ±10% is often used, where this threshold is used as a measure of the quality of the calculated ages. in contrast to agedisplay, the complete data set including the ‘discordant’ data is not shown behind the ‘concordant’ data, but separately using the same axis as for the concordant data. • use fixed bin size (ma/bin) allows a histogram of filtered ages to be plotted (based on the concordant data obtained via the concordia filter), where the bin size in million years (ma) is manually chosen. if this checkbox is deselected, a calculated value for the bin size is used, based on the formulae in doane (1985). • display max probability (checkbox) shows the ages of the calculated probability peaks for the ‘concordant’ data set. • min/max age controls the age range to be shown in the chart. additional parameters can be set through right-clicking on the produced chart, where the title, chart title, axis range, appearance and colours can be modified. jagedisplay can display a single chart for one data set or stacked charts of several data sets (fig. 2). stacked charts are useful for comparison of data sets sampled, e.g. at different geographical, stratigraphical or lithological locations in order to identify trends or differences in the age distributions. the stacking follows the order chosen in the selected file items (or selected db items) window at the right, and can be rearranged as appropriate. as the purpose of stacking is to compare, the stacked data sets are displayed for the same age range as set by the min/max age control. use of jagedisplay in sedimentary provenance and petrological studies detrital studies typically aim to characterise the age population(s) from a sample by means of a large number of single grain analyses, e.g. obtained by la-icpms. in statistical terms the goal of such studies is to estimate the so-called probability density function that gives the relative likelihood of the different ages in the population (vermeesch 2012). zircon is the most common mineral used for detrital studies, but other mineral phases like rutile and titanite are increasingly being employed in order to obtain supplementary information (e.g. zack et al. 2004; stendal et al. 2006; meinhold et al. 2008; thomsen et al. 2015; 105 bruand et al. 2016). signatures of sediment sources not represented in the zircon data set may be present in detrital titanite, rutile or apatite populations. in addition, age data on these minerals may provide further chronological and petrogenetic insight into the tectono-thermal history of their source regions (e.g. mcateer et al. 2010, 2014; knudsen et al. 2015; thomsen et al. 2015). similarily, in metamorphic and magmatic petrology, complex zoning patterns that correspond to distinct geological events can be recorded in zircon, baddeleyite, titanite, rutile, apatite, monazite, xenotime, allanite or any other accessory mineral phase for which the age can be determined by u-th/ pb methods and can be assessed in probability density plots diagrams through the jagedisplay software. future options in jagedisplay although pdps constitute the most widespread method used today for displaying detrital age distributions, they lack a firm theoretical basis as probability density estimators, and this may produce counter-intuitive results when the number of analyses and/or their quality (precision) is high (vermeesch 2012). accordingly, an alternative and robust standard statistical technique, the kernel density estimation (kde) as described in vermeesch (2012), will be included in jagedisplay, so that users can compare the two methods. we also expect to include a more versatile option for age distribution plots, where any two isotopic pairs can be used in a pdp or kde diagram, in a similar way as the built-in option for combined 206pb/238u and 207pb/206pb age distributions. fig. 2. stacked chart of combined histograms and pdp charts from 1100 to 4000 ma of zircons from seven samples using 207pb/206pb ages, a fixed bin size of 50 ma/bin and a concordia filter of ±10%. 106106 detailed statistics reporting information on the calculated maximum probability peaks can be useful for improved age differentiation of, e.g. magmatic or metamorphic episodes. at present, this information is only indirectly reported through the plot axis. we expect to include this information in pop-up windows for single peaks and as a print-out option including the statistics of all peaks or a chosen age section of the plot. finally, the jagedisplay software offers an option to load input files from a database through a local server connection (file => connect to database). at present, this option only allows reading of certain data types from geus’ oracle server. it is our intention to expand this and thus provide users with access to data from a local server repository. additional remark in this presentation only u/pb geochronological data obtained by la-icpms analysis on zircon are used. jagedisplay can of course also evaluate u-th-pb data acquired by other instrumentation and complex age distribution patterns measured from other accessory phases such as titanite, rutile, apatite, monazite, allanite, perovskite, xenotime or other minerals that can incorporate u or th into their crystal structure. references bruand, e., storey, c. & fowler, m. 2016: an apatite for progress: inclusions in zircon and titanite constrain petrogenesis and provenance. geology 44(2), 91–94, http://dx.doi.org/10.1130/g37301.1 doane, d.p. 1985: aesthetic frequency classifications. the american statistician 30, 181–183. fergusson, c.l., fanning, c.m. & green, t.j. 2001: proterozoic–cambrian detrital zircon and monazite ages from the anakie inlier, central queensland: grenville and pacific-gondwana signatures. australian journal of earth sciences 48(6), 857–866. knudsen c., thomsen t.b. & hinchey, a. 2015: detrital zircon, rutile and titanite investigations from present-day labrador river drainages: fingerprinting the grenvillean front. goldschmidt conference, prague, 16–21 august 2015. abstract 1627 only. ludwig, k. 2003: user’s manual for isoplot 3.00: a geochronological toolkit for microsoft excel. berkeley geochronology center special publication 4, 74 pp. mcateer, c.a., daly, j.s., flowerdew, m.j., connelly, j.n., housh, t.b. & whitehouse, m.j. 2010: detrital zircon, detrital titanite and igneous clast u–pb geochronology and basement–cover relationships of the colonsay group, sw scotland: laurentian provenance and correlation with the neoproterozoic dalradian supergroup. precambrian research 181, 21–42. mcateer, c.a., daly, j.s., flowerdew, m.j., whitehouse, m.j. & monaghan, n.m. 2014: sedimentary provenance, age and possible correlation of the iona group, sw scotland. scottish journal of geology 50, 143–158. meinhold, g., anders, b., kostopoulos, d. & reischmann, t. 2008: rutile chemistry and thermometry as provenance indicator: an example from chios island, greece. sedimentary geology 203, 98–11. morton, a.c., claoué-long, j.c. & berge, c. 1996: shrimp constraints on sediment provenance and transport history in the mesozoic statfjord formation, north sea. journal of the geological society, london 153, 915–929. paton, c., hellstrom, j.c., paul, p., woodhead, j.d. & hergt, j.m. 2011: iolite: freeware for the visualisation and processing of mass spectrometric data. journal of analytical atomic spectrometry 26(26), 2508–2518. pell, s.d., williams, i.s. & chivas, a.r. 1997: the use of protolith zircon-age fingerprints in determining the protosource areas for some australian dunes sands. sedimentary geology 109, 233–260. rainbird, r.h., mcnicoll, v.j., theriault, r.j., heaman, l.m., abbott, j.g., long, d.g.f. & thorkelson, d.j. 1997: pan-continental river system draining grenville orogen recorded by u-pb and smnd geochronology of neoproterozoic quartzarenites and mudrocks, northwestern canada. journal of geology 105, 1–17. sircombe, k.n. 1999: tracing provenance through the isotope ages of littoral and sedimentary detrital zircon, eastern australia. sedimentary geology 124, 47–67. sircombe, k.n. 2004: agedisplay: an excel workbook to evaluate and display univariate geochronological data using binned frequency histograms and probability density distributions. computers and geosciences 30, 21–31. stendal, h., toteu, s.f., frei, r., penaye, j., njel, u.o., bassahak, j., nni, j., kankeu, b., ngako, v. & hell, j.v. 2006: derivation of detrital rutile in the yaoundé region from the neoproterozoic pan-african belt in southern cameroon (central africa). journal of african earth sciences 4(4), 443–458. thomsen, t.b., knudsen, c. & hinchey, a. 2015: detrital zircon, rutile and titanite investigations from present-day labrador river drainages: fingerprinting the grenvillean front. geological survey of denmark and greenland bulletin 33, 77 –80. vermeesch, p. 2012: on the visualisation of detrital age distributions. chemical geology 312–313, 190–194, http://dx.doi.org/10.1016/j. chemgeo.2012.04.021 zack, t., von eynatten, h. & kronz, a. 2004: rutile geochemistry and its potential use in quantitative provenance studies. sedimentary geology 171(1), 37–58. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: tbt@geus.dk geological survey of denmark and greenland bulletin 17, 2009, 53-56 in recent years, both the petroleum industry and government research institutions have shown renewed interest in the petroleum potential of the high arctic. at the same time, a range of activities are taking place, aimed at defining national borders in the arctic ocean following ratification of article 76 of the united nations convention on the law of the sea (unclos). parallel to the general upsurge in data acquisition activities, the united states geological survey has carried out a circum-arctic resource appraisal (cara), which for north-east greenland was published in 2007. this assessment indicated that a significant petroleum exploration potential exists on the north-east greenland shelf, in particular in the danmarkshavn basin and the north danmarks havn salt province (fig. 1). the estimated potential amounts to 31 billion barrels of oil equivalents, principally in the form of natural gas. for comparison, this roughly corresponds to one third of the original reserve of the north sea basins. the geology of the danmarkshavn basin and offshore areas farther to the north is only known in broad outline, since no wells have been drilled and only reconnaissance geophysical data are available. moreover, the extensive ice cover and the overall hostile climate of the region pose significant logistical and technical challenges to data acquisition. clearly, this emphasises the importance of analogue studies based on the much better known geology of the onshore basins in east and north-east greenland. in 2007/2008, the geological survey of denmark and greenland (geus) launched a major petroleum industry-sponsored project with the objective of updating and expanding our current understanding of the petro leum geology of east and north-east greenland. the project is planned to continue for the next four to five years, and includes compilation of relevant existing data in the form of a geographic information system (gis) product, supplemented by new data obtained from shallow core drilling and new field work. below we give a brief overview of a range of field activities that took place in east and north-east green land in the summer of 2008. shallow core drilling in jameson land the objective of the drilling was to recover core material from the upper jurassic organic-rich shale of the katedralen member (hareelv formation) to obtain a reference section of shallow core drilling and petroleum geology related field work in east and north-east greenland 2008 jørgen a. bojesen-koefoed, morten bjerager and stefan piasecki © geus, 2009. geological survey of denmark and greenland bulletin 17, 53–56. available at: www.geus.dk/publications/bull c c' b' a a' 74°n 76°n 72°n 70°n lle saf pdmf df lle cretaceous jurassic triassic permian fault buried deepseated faults stauning alper fault post-devonian main fault dombjerg fault liverpool land escarpment drill site near blokelv tertiary milne land traill ø gauss halvø ymer ø geographical society ø hold with hope home forland clavering ø gael hamke bugt th.thomsen land kuhn ø daneborg rødryggen brorson halvø hochstetter forland store koldewey wollaston forland pd m f sa f d f kap dalton 100 km 26°w28°w 26°w 24°w jameson land liverpool land mestersvig schuchert dal fortet constable pynt greenland 18°w 16°w22°w 20°w blo sse vil le kyst h ud so n la nd ndsp ndsp: north danmarkshavn salt province danmarkshavn basin ugleelv kejser franz joseph fjord kong oscar fjord ca rlsb erg fj ord hall bredning scoresby sund hurry inlet fig. 1. map of east and north-east greenland, showing the distribution of permian to neogene sedimentary rocks and the location of place names mentioned in the text. ndsp: north danmarkshavn salt basin. 53 rosa_2008:rosa-2008 01/07/09 15:48 side 53 the most prolific source rock in the north atlantic region (fig. 2). the deepest and most central part of the jameson land basin at blokelv approximately 35 km west of con stable pynt was chosen as the drill site (fig. 1). the blokelv drilling (fig. 3) produced high quality cores (ggu 511101) with a diameter of 5.6 cm and 99.3% recovery to a depth of 233.8 m (fig. 4). the top part from 1.72 to 10.08 m recovered homogeneous medium to fine-grained sandstone with two thin shale intervals of the sjællandselv member (hareelv formation). the katedralen member (hare elv formation) from 10.08 to 233.8 m consists of alternating sandstone and shale units; the lower boundary of the member was not reached. studies of the new cores from the katedralen member focus on the stratigraphy and age of the hareelv formation and lateral, contemporaneous deposits (fig. 2). the lower boundary of the katedralen member is well known from the ugleelv region and the west side of hurry inlet, where few, condensed beds of mudstone and finegrained sandstone separate the fossil bjerget and hareelv formations. these beds are age equivalent to the olympen formation in central to northern jame son land (p. athleta to c. densiplicatum chronozones of late callovian to middle oxfordian age; larsen & surlyk 2003). the deposition of laminated mudstone of the katedralen member began in the late oxfordian. the continued sedimentation of the katedralen member mud reaches into the kimmeridgian and volgian especially in eastern and southern jameson land (fig. 2). during the volgian, the coarsegrained sand of the sjællandselv member and raukelv for mation was deposited in the western and central parts of the jameson land basin at the same time as the katedralen member mud was deposited in the easternmost, apparently much deeper basin (surlyk 2003). the youngest parts of the katedralen member are only known from the locality fortet near constable pynt in eastern jameson land, where the dinoflagellate assemblage indicates an age not older than latest middle volgian (fig. 2). the upper boundary of the kate dralen member is located where laminated mudstone is replaced by homo geneous sandstone of the sjæl landselv member. the sand slumped into the central basin from shelf-edge deltas prograding from north-west and west. this shift in deposition can be followed from west to east across the basin in a number of shallow core drillings and exposed 54 chronostratigraphy age (ma) stage volgian berriasian tithonian kimmeridgian oxfordian callovian jameson land c re t. ju ra ss ic lo w . u pp er m id dl e raukelv fm zeus mb no data athene mb goniomyakløft mb fossilbjerget fm condensed katedralen mb hades mb sjællandselv mb parnas mb s n u m l u l u m l u m l 160 150 shallow marine sandstones shelf transition-zone siltstones/heteroliths marine shelf/basinal mudstones deep marine sandstones hiatus/condensed mass flow sandstone source rock h ar ee lv f m fig. 2. chronoand lithostratigraphy of the middle and upper jurassic in jameson land. modified from surlyk (2003). fig 3. drilling operations near blokelv in jameson land. three-metre-long cores are drilled at a time, and each time a core is retrieved water is pouring out of the drill hole. rosa_2008:rosa-2008 01/07/09 15:48 side 54 sections. the boundary is strongly diachronous: kimmerid gian–volgian in the west to middle – upper volgian in the easternmost part of the basin. an extensive analytical programme to evaluate source and reservoir rock properties to establish a high-resolution dino flagellate cyst and macrofossil biostratigraphic zonation, and assess sedimentological, diagenetic and sequence stratigra phic aspects of the penetrated succession is in progress. field work field work took place by means of helicopter-supported field teams of 2–4 persons, operating primarily out of field camps. the activities covered a wide range of geological disciplines that are briefly described below. field teams camped at or visited locations from kap dalton on the blosseville kyst in the south to kuhn ø in the north (fig. 1). volcanology volcanic rocks in the form of plateau basalts as well as various intrusions abound in east and north-east greenland, and their histories of emplacement and mutual relationships are important elements in the understanding of the regional geology. a field team sampled and studied the youngest volcanic rocks in the region that are preserved in a small graben at kap dalton. additional studies were undertaken on hold with hope and wollaston forland and on nearby islands east and north thereof. samples collected from various outcrops will be subjected to chemical and petrological analyses as well as radiometric dating in order to establish their ages and mutual relationships. petroleum source rocks and cretaceous– palaeogene stratigraphy petroleum source rocks are known to be present in east and north-east greenland in a number of stratigraphic intervals ranging in age from middle devonian to upper jurassic, whereas source rocks in younger units still remain to be demonstrated. among these various units the upper jurassic ‘kimmeridge clay equivalents’ referred to as the hareelv, kap leslie and bernbjerg formations (surlyk 2003) must be considered the more important ones with respect to petroleum exploration. the documentation of spatial and temporal variations in petroleum potential within these units is crucial to the prediction of the distribution of the potential in the little-known offshore basins. upper jurassic shale of the bernbjerg formation was sampled on hold with hope and wollaston forland, and potential drill sites were found at rødryggen and brorson halvø on wollaston forland. in addition, cretaceous shale was sampled for both stratigraphic and geochemical purposes on wollaston forland, whereas sampling of cretaceous deposits and drill-site identification on hold with hope had to be postponed due to the presence of polar bears in the area. the stratigraphy and sedimentological development of the cretaceous–palaeogene succession of north-east green land have hitherto received relatively little attention, probably due to a general assumption that it was deposited during a period of tectonic quiescence with little change in overall basin configuration. during the field season in 2008, some effort was devoted to increasing our knowledge of the cretaceous–palaeogene development through sedimentological studies and stratigraphic sampling. among the results was the discovery of a more than 150 m thick, unmapped succession of presumably palaeogene sand with minor mud55 fig. 4. core box no. 62 with black organic rich mudstone (m), intrusive sandstone dykes (is), and laminated heterolithic sandstone-mudstone (ls-m) of the katedralen member, hareelv formation. the diameter of the core is 56 mm. top 229.66 m 0.1 m bottom 233.40 m ls-m is m rosa_2008:rosa-2008 01/07/09 15:48 side 55 56 authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: jbk@geus.dk stones and coal directly underlying the plateau basalts on eastern wollaston forland. sub-basaltic palaeogene sand is reported from several localities in the region. landscape analysis, uplift and sand provenance studies burial and uplift are geological processes, the magnitude and timing of which are of utmost importance in petroleum geology. assessment of these may be undertaken by a combination of methods, including large-scale landform analysis and apatite fission track analysis, which enable us to detect ancient erosion or peneplanisation surfaces as well as quantify palaeogeothermal gradients (bonow et al. 2007). moreover, identification of the source areas for sandy deposits through geological time is important to understand temporal dispersal patterns of sediment within the basin, and for correlation of sand units. such provenance studies can be carried out by age determinations and detailed chemical analyses of provenance-sensitive minerals such as zircon and garnet. since up lift and provenance studies to a large extent make use of the same type of sample material, two field teams collected samples for both purposes. spot and profile samples were collect ed at a large number of locations from milne land and jame son land in the south to wollaston forland in the north. seepage studies experience from central west greenland shows that when certain requirements are met petroleum seepages may be preserved in the lower parts of the palaeogene basalt succession and in carbonate veins associated with dykes (bojesenkoefoed et al. 1999). the utility of the experience gained in west greenland was tested in north-east greenland, but with little success, since essentially no traces of petroleum were found in the volcanic rocks exposed in the region. a number of locations on kuhn ø and wollaston forland were checked, but no hydrocarbons were found. the lavas are indeed very porous, but the level of thermal alteration indicated by the zeolite facies is too low, the volcanic rocks are too thin and healed carbonate-filled veins are nearly absent. similar studies of the volcanic rocks on hold with hope were severely hampered by bad weather and polar bears, but since the geological conditions appear more favourable there and seepages have been recorded previously, further attempts should be made during future field work. airborne stereo-photography systematic stereo-photography is a valuable tool for largescale stratigraphy and structural geology (dueholm & pedersen 1992). nearly 2000 stereo-photographs were taken from a partenavia aircraft during two sessions in (1) jameson land, milne land, schuchert dal, kong oscar fjord, southern traill ø region and (2) kejser franz joseph fjord, gauss halvø, hudson land, gael hamke bugt, home forland and the south-eastern hold with hope region. future activities in north-east greenland the project described here is planned to continue for the next 4–5 years, with both field work and shallow core drilling starting in the south and gradually moving northwards. in addition, an excursion to the region for sponsoring oil companies is planned for 2010. in 2008, the bureau of minerals and petroleum, greenland, published a ‘roadmap’ for a future licensing round in north-east greenland which, pending political approval, will lead to nomination and licensing rounds in 2011–2013. in order to meet the needs of the sponsors, further special studies may be undertaken in the forthcoming years. acknowledgements the staff at the sirius sledge patrol stations at daneborg and mestersvig and the staff at constable pynt airport are thanked for their practical assistance and hospitality. references bojesen-koefoed, j.a., christiansen, f.g., nytoft, h.p. & pedersen, a.k 1999: oil seepage onshore west greenland: evidence of multiple source rocks and oil mixing. in: fleet, a.j. & boldy, s.a.r. (eds): petroleum geology of northwest europe: proceedings of the 5th conference, 305–314. london: geological society. bonow, j.m., japsen, p., green, p.f., wilson, r.w., chalmers, j.a., klint, k.e.s., van gool, j.a.m., lidmar-bergström, l. & pedersen, a.k. 2007: a multi-disciplinary study of phanerozoic landscape development in west greenland. geological survey of denmark and greenland bulletin 13, 33–36. dueholm, k.s. & pedersen, a.k. (eds) 1992: geological analysis and mapping using multi-model photogrammetry. rapport grønlands geologiske undersøgelse 156, 72 pp. larsen, m. & surlyk, f. 2003: shelf-edge delta and slope deposition in the upper callovian – middle oxfordian olympen formation, east green land. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 931–948. surlyk, f. 2003: the jurassic of east greenland: a sedimentary record of thermal subsidence, onset and culmination of rifting. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 659–722. rosa_2008:rosa-2008 01/07/09 15:48 side 56 geological survey of denmark and greenland bulletin 20, 2010, 79–82 79 greenland is receiving unprecedented international attention, both in scientific and political circles. characterised by a central ice sheet up to 3.4 km thick (inland ice), numerous ice caps and hundreds of outlet glaciers debouching into the surrounding oceans, greenland supports the second largest ice mass in the world. analysis of glacier movements, melt rates and ice loss to the sea, provide data with which to assess mass balance changes and thereby predict global sealevel rise. thus greenland plays a central role in the current worldwide debate on climate change. present-day dynamic ice loss is invariably advertised by the fast moving glaciers of western greenland with their spectacular calf ice production, such as the ice streams around disko bugt reviewed by weidick & bennike (2007). this tends to overshadow ice stability and expansion seen in the form of stationary and advancing glaciers elsewhere in greenland (modis 2009). while the seawards acceleration of glacier flow and retreat in frontal positions can be readily attributed to a shift in atmospheric and oceanic conditions (global warming), the same explanation can hardly be used for glaciers with contrasting movement histories. aim of this paper we focus on three west coast, marine-terminating glaciers between 75° and 78°n (fig. 1) to elucidate the retreat–advance paradox referred to above. steenstrup gletscher and tracy gletscher are chosen to illustrate the regional pattern of ice recession, including massive ice wasting on a broad front in melville bugt, whereas berlingske bræ defies this trend by long-lasting advance. the receding glaciers steenstrup and tracy are known from regional surveys (e.g. kollmeyer 1980; rignot & kanagaratnam 2006), but the advancing glacier berlingske bræ has only been cursorily mentioned in map descriptions (dawes 1992, 2006). we present maps showing the terminus fluctuations of the three glaciers based on historical records (figs 2–4) but the paper’s four-page limit prohibits discussion of the early sources. this aspect, climatic records and their relation to ice fluctuations, and comparisons with other greenland glaciers, will be dealt with in a forthcoming paper. historical sources and the 2009 database t.c. chamberlin and r.d. salisbury were the first to investigate glaciers in the region in 1894–95 when they reached as far north as inglefield bredning, the location of our northern glacier (tracy gletscher). the next milestone was the regional mapping by geologist and cartographer lauge koch between 1916 and 1923 who surveyed and described glaciers throughout the region. the 1940s heralded a new era of research with the incoming of aerial photographs and such images are available from the period 1948–1985. finally in the last decades, satellite images have assured a continual record of the areas covered and uncovered by greenland glaciers (weidick 1994). in this paper we make use of such imagery from the period 1963–2009. table 1 summarises our data sources. an advancing glacier in a recessive ice regime: berlingske bræ, north-west greenland peter r. dawes and dirk van as fig. 1. map of north-west greenland showing the locations of the three studied glaciers featured in figs 2–4. © geus, 2010. geological survey of denmark and greenland bulletin 20, 79–82. open access: www.geus.dk/publications/bull greenland 60°w 70°w 76°n 77°n 100 km fig. 4 tracy gletscher fig. 3 berlingske bræ fig. 2 steenstrup gletscher harald moltke bræ prudhoe land 78°n inland ice melville bugt baffin bay wolstenholme fjord l a u g e k o c h k y s t steensby land pituffik (thule air base) qaanaaq kap york 8080 steenstrup gletscher, melville bugt steenstrup gletscher is the widest glacier of the impressive ice front that characterises lauge koch kyst and that calves into melville bugt (figs 1, 2). stretching from the kap seddon peninsula to red head, where steenstrup gletscher borders the fairly stable kjer gletscher, the glacier has an irregular and crevassed floating tongue more than 30 km wide. ice production from the central part is spectacular, both as regards the number of calved bergs and their size. koch (1928) noted that large portions of the 25 m high floating tongue become detached and move seawards before being broken up, while kollmeyer (1980) described the calving of ‘jigsaw puzzle type’ icebergs in excess of 1 km in length. as illustrated by fig. 2, the main fluctuations affected the central part of the floating tongue where drawback since 1916 is almost 25 km. melville bugt is noted for its ice-infested waters during summer months and the records show that the nature and position of the ice front can change seasonally depending on the degree of ice congestion. in five months, from spring to autumn 1916, the central part of the front had moved westwards by c. 1 km while in 1920, after a summer when melville bugt was free of ice, there was recession of more than 6 km (koch 1928). in contrast, the northern segment of the terminus was stationary between 1916 and 1923. fig. 2. satellite image of steenstrup gletscher and neighbouring glaciers, southern melville bugt, showing eight frontal positions from 1916 to 2009. thin, white line: coastline. for sources, see table 1. fig. 3. satellite image of berlingske bræ, granville fjord showing seven frontal positions from 1922 to 2009. thin, white line: coastline. for sources, see table 1. table 1. data for north-west greenland glacier fluctuations year type/medium source 1892 field/map peary (1892) 1916 field/map koch (1922) 1922 field/map koch (1932) 1949* oap†/map geodetic institute§ 1953 vap‡ u.s. military 1963 satellite zhou & jezek (2003) 1971 vap‡ greenarctic consortium 1985 vap‡ geodetic institute 1975 1999 landsat nasa and u.s. 2007 satellite geological survey 2009 *photography 1948–1950; §1st edition map 1954 †oblique aerial photography; ‡ vertical aerial photography red head kap seddon 10 km sverdrup gletscher diet ric hs on g l. steenstrup gletscher kjer gletscher 1916 1949 1963 1975 1985 1999 2007 2009 5 km 1916 1949 gra nv ille f jor d berlingske bræ pol itik en bræ 1963 1975 1985 2007 2009 81 the shrinkage shown in fig. 2 has produced new bedrock exposures, and red head, which was a semi-nunatak in 1916, became an island in 2005. berlingske bræ, granville fjord berlingske bræ flows westwards into the head of granville fjord from the major ice cap of steensby land that has a bridge connection with the inland ice (figs 1, 3). it is the only glacier of several draining into granville fjord that currently reaches the sea. the glacier has a rather low gradient and its flow pattern can be traced for about 25 km before being lost to the ice cap. the glacier trunk is 3–4 km wide, and today the terminus is irregular, rather slender and crevassed, and over 2 km across. it is unknown whether the snout is afloat, but the apparent lack of iceberg production suggests it is grounded. an unnamed tributary originating from a more westerly ice cap joins the northern flank of berlingske bræ contributing to its westerly flow into granville fjord. the terminus positions shown in fig. 3 illustrate a continuous ice advance in excess of 4 km in the last 85 years that has changed the glacier front from being terrestrial to marine. as opposed to when the glacier terminated on land, its tongue now is strongly tapering. the glacier has overrun the entire alluvial gravel plain that in the early 20th century extended beyond its snout and separated it from the fjord, and it has also engulfed bedrock exposures on its southern flank that were mapped by the first author in 1974. tracy gletscher, inglefield bredning tracy gletscher is the second largest of six outlet glaciers that debouch into the headwaters of inglefield bredning (figs 1, 4). it is about 5 km wide with a steep front that has a regular concave trace which is probably afloat. the flow pattern of the glacier is recognisable over 30 km before being lost to the east to the inland ice proper. the 115-year record (fig. 4) shows substantial ice wasting amounting to frontal recession of c. 15 km. the terminus positions show that for a century (1892–1985) tracy gletscher was coalesced with farquhar gletscher. in 1923 the seaward front of the floating confluent ice embracing melville gletscher was a cliff 20 m high and 19 km long (koch 1928). over the past century the glacier tongue has lost about 100 km2 of ice, which represents at least 20 km3 based on koch’s observation. recession with the impressive break-up of the confluent ice mass has led to striking landscape changes. for example, longtime nunataks lee bjerg and field bjerg are now lapped by the sea, while josephine peary ø finally lives up to its name as an island being eventually released from the ice around 1960 (inûterssuaq uvdloriaq, personal communication 1971). recent glacial history: retreat versus advance the glaciers of north-west greenland and their marginal deposits are shown on the thule 1:500 000 scale geological sheet, and a summary of glacial history is given in the map description (dawes 2006). the regional pattern of spatial change seen in terms of terminal positions between 1948–50 and 1985 is shown on the maps of dawes (1988). this information, plus the early records summarised by koch (1928), the regional analysis of davies & krinsley (1962) and satellite data of the last decades, demonstrate that the general recession of the inland ice and its outlet glaciers is regional in character and persistent for more than a century. the drastic deglaciation of melville bugt has brought the ice limit there close to the early holocene position (bennike 2008). harvard øer lee bjerg tracy gletscher smithson bjerge heilprin gletscher inglefield bredning josephine peary ø m el vi lle g l. field bjerg sh ar p g le ts ch er fa rq uh ar g le ts ch er 5 km 1892 1922 1949 1963 1975 1985 2007 2009 fig. 4. satellite image of tracy gletscher and neighbouring glaciers, inglefield bredning, showing eight frontal positions from 1892 to 2009. thin, white line: coastline. for sources, see table 1. 8282 the overall pattern is that glaciers with floating tongues, like steenstrup and tracy, have shown by far the largest retreat and ice wastage. the most extensive ice withdrawal has been along the heavily glaciated lauge koch kyst, where the lowering ice surface is being pierced by its rock substratum and where nunataks have become shoreline, and ice-rooted peninsulas insular. however, some glaciers show current fast retreat after decades of stability (e.g. sverdrup gletscher; fig. 2). in general, land-based glaciers show relatively sluggish movement, and some have been almost stationary or only show minor retreat (e.g. prudhoe land glaciers; dawes 2006). it is clear that given the existence of a detailed database, the general recession can be seen to have been interrupted by short periods of comparative stability and even advance (e.g. harald moltke bræ; mock 1966). in contrast to this regional recessive regime, berlingske bræ shows continual advance for at least 85 years. an explanation of this deviant behaviour must be sought in the fact that the glacier originates from an independent ice cap that responds to changes in temperature and precipitation differently than the inland ice. the glacier advance can be a response to increased precipitation on the ice cap or increased basal sliding, both of which could be related to the observed increase in atmospheric temperatures. no matter which, the advance of berlingske bræ over such a long period is unexpected in a warming climate. conclusions, relevance to global climate research and future work berlingske bræ is located between steenstrup gletscher and tracy gletscher that are 340 km apart. the two receding glaciers compare with others in melville bugt (and in other areas of greenland) indicating changed mass balance of their source: the inland ice. the main causes of this long-lasting change – documented in our data back to 1892 – must be regional, and thus the present warming climate must affect the process. however, whatever the fundamental cause (or causes) controlling regional meltdown, it has been a subordinate factor at berlingske bræ where there is long-standing advance. seen in terms of the regional, recessive ice regime in which it is located, berlingske bræ is anomalous and thus outside mainstream research concerning analysis of dwindling ice masses and their response to global warming. however, if we are to understand the underlying complex processes, and ultimately the effect of climate change on the regional recessive regime, attention should also be paid to such glaciers. among other things, this research should be directed to discovering why receding and expanding glaciers with century-long contrasting histories occur side by side. this paper is a contribution to international promotion of this aspect of glacioclimatic research in progress at the survey, both in our study region and elsewhere in greenland (e.g. weidick 2009). references bennike, o. 2008: an early holocene greenland whale from melville bugt, greenland. quaternary research 69, 72–76. davies, w.e. & krinsley, d.b. 1962: the recent regimen of the ice cap margin in north greenland. international association of scientific hydrology 58, 119–130. dawes, p.r. 1988: geological map of the thule district, north-west greenland. 1:100 000 sheets 1–6 and 1:200 000 sheets 7–11. unpublished maps, geological survey of denmark and greenland, copenhagen. dawes, p.r. 1992: new geological map of the thule region, north-west greenland. rapport grønlands geologiske undersøgelse 155, 42–47. dawes, p.r. 2006: explanatory notes to the geological map of greenland, 1:500 000, thule, sheet 5. geological survey of denmark and greenland map series 2, 97 pp. koch, l. 1922: note to maps of melville bay from wilcox point to cape york and of north greenland from 81°–83°35́ n, 38°–56°w. meddelelser om grønland 64(2), 77–88. koch, l. 1928: contributions to the glaciology of north greenland. meddelelser om grønland 65(2), 183–464. koch, l. 1932: map of north greenland, scale 1:300,000. copenhagen: geodetic institute, 19 sheets. kollmeyer, r.c. 1980: west greenland outlet glaciers: an inventory of the major iceberg producers. cold regions science and technology 1, 175–181. mock, s.j. 1966: fluctuations of the terminus of the harald moltke bræ, greenland. journal of glaciology 6(45), 369–373. modis 2009: modis studies of greenland. moderate resolution imagery spectroradiometer, nasa and byrd polar research center, http//bprc.osu.edu/modis/?p=61. peary, r.e. 1892: the north greenland expedition of 1891–92. journal of american geographical society 24, 536–558. rignot, e. & kanagaratnam, p. 2006: changes in the velocity structure of the greenland ice sheet. science 311, 986–990. weidick, a. 1994: satellite image atlas of glaciers of the world. greenland. united states geological survey professional paper 1386–c, 141 pp. weidick, a. 2009: johan dahl land, south greenland: the end of a 20th century glacier expansion. polar record 45(235), 337–350. weidick, a. & bennike, o. 2007: glaciation history and glaciology of jakobshavn isbræ and the disko bugt region, west greenland: a review. geological survey of denmark and greenland bulletin 14, 78 pp. zhou, g. & jezek, k. 2003: disp yearly satellite photographic mosaics of greenland 1962–1963. national snow and ice data center, boulder, colorado. digital media. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: prd@geus.dk geological survey of denmark and greenland bulletin 11, 179-184 179 magnetic anomalies and metamorphic boundaries in the southern nagssugtoqidian orogen, west greenland john a. korstgård, bo møller stensgaard and thorkild m. rasmussen within the southern nagssugtoqidian orogen in west greenland metamorphic terrains of both archaean and palaeoproterozoic ages occur with metamorphic grade varying from low amphibolite facies to granulite facies. the determination of the relative ages of the different metamorphic terrains is greatly aided by the intrusion of the 2 ga kangâmiut dyke swarm along a nne trend. in archaean areas dykes cross-cut gneiss structures, and the host gneisses are in amphibolite to granulite facies. along itilleq strong shearing in an e–w-oriented zone caused retrogression of surrounding gneisses to low amphibolite facies. within this itivdleq shear zone kangâmiut dykes follow the e–w shear fabrics giving the impression that dykes were reoriented by the shearing. however, the dykes remain largely undeformed and unmetamorphosed, indicating that the shear zone was established prior to dyke emplacement and that the orientation of the dykes here was governed by the shear fabric. metamorphism and deformation north of itilleq involve both dykes and host gneisses, and the metamorphic grade is amphibolite facies increasing to granulite facies at the northern boundary of the southern nagssugtoqidian orogen. here a zone of strong deformation, the ikertôq thrust zone, coincides roughly with the amphibolite–granulite facies transition. total magnetic field intensity anomalies from aeromagnetic data coincide spectacularly with metamorphic boundaries and reflect changes in content of the magnetic minerals at facies transitions. even the nature of facies transitions is apparent. static metamorphic boundaries are gradual whereas dynamic boundaries along deformation zones are abrupt. keywords: aeromagnetic data, magnetic anomalies, metamorphic facies, nagssugtoqidian orogen, west greenland __________________________________________________________________________________________________________________________________________________________ j.a.k., department of earth sciences, university of aarhus, høegh-guldbergsgade 2, dk-8000 århus c, denmark. e-mail: john.korstgard@geo.au.dk b.m.s. & t.m.r., geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350 copenhagen k, denmark. the establishment of the palaeoproterozoic nagssugtoqidian orogen in west greenland (ramberg 1949) is based on the deformation and metamorphism of the kangâmiut dykes, dated at 2.04 ga by nutman et al. (1999). south of the southern nagssugtoqidian front (snf in fig. 1), in the southern nagssugtoqidian foreland, kangâmiut dykes are undeformed and cross-cut gneiss structures. north of the front, gneisses and dykes have been metamorphosed and deformed together during the nagssugtoqidian orogeny. here, gneiss structures and dyke margins are concordant and dykes transformed into amphibolites. this is the simple story upon which ramberg (1949) based his definition of the ‘nagssugtoqides’. ramberg also divided the nagssugtoqidian orogen into three metamorphic complexes based on the metamorphic grade of the rocks. thus the egedesminde complex was the northernmost amphibolite facies complex, the isortoq complex the central granulite facies complex, and the ikertôq complex the southernmost amphibolite facies complex. the current division of the orogen (fig. 1) is based on structural criteria, and division boundaries now follow major structural features (marker et al. 1995). the current division therefore deviates considerably from ramberg’s original division for the northern and central nagssugtoqidian orogen, whereas the © geus, 2006. geological survey of denmark and greenland bulletin 11, 179–184. available at: www.geus.dk/publications/bull 180 50 km inland ice ? ? snf cno nno aasiaataasiaat kangerlussuaqkangerlussuaq sno nordre strømfjord shear zone nordre isortoq steep belt southern nagssugtoqidian foreland sø nd re str øm fjo rd nordre strømfjord aasiaat kangerlussuaq nordre strømfjord arfersiorfik fig. 3 fig. 2 sisimiut ikertôq thrust zone itivdleq shear zone 54° 67° 68° 51° surficial deposits quaternary north atlantic craton granodioritic-granitic gneiss (northern parts reworked) orthogneiss (largely unreworked) metasedimentary rocks (palaeoproterozoic, may include archaean components) nagssugtoqidian orogen sisimiut charnockite (palaeoproterozoic) arfersiorfik quartz diorite (palaeoproterozoic) orthogneiss (archaean, reworked) metasedimentary rocks (archaean, may include proterozoic components) amphibolite anorthosite and ultrabasic rocks 500 km greenland fig. 1. schematic geology of the southern part of the nagssugtoqidian orogen and adjacent forelands (modified from escher & pulvertaft 1995 and marker et al. 1995). sno, southern nagssugtoqidian orogen; cno, central nagssugtoqidian orogen; nno, northern nagssugtoqidian orogen; snf, southern nagssugtoqidian front. the locations of thrust and shear zones are defined from trends observed in the aeromagnetic data; note that the e–wtrending thrust zone with question marks north of kangerlussuaq is uncertain, as this structure has not been confirmed by geological mapping. black frames show the locations of figs 2, 3. southern nagssugtoqidian orogen corresponds almost exactly to ramberg’s original ikertôq complex. the southern nagssugtoqidian orogen (sno in fig. 1) in the coastal region between sisimiut and itilleq consists mainly of quartzofeldspathic gneisses of granodioritic to tonalitic composition. several supracrustal layers occur, particularly in the northern part of the sno. the supracrustal rocks are mainly garnet-biotite schists, rusty weathering biotite gneisses and amphibolites. the metamorphic grade is low amphibolite facies to granulite facies, and due to the fortunate timing of the intrusion of the kangâmiut dykes it is possible to assign relative ages to the different metamorphic terrains in the region. pre-dyke metamorphism and deformation south of and immediately north of itilleq, the kangâmiut dykes are largely undeformed, unmetamorphosed and 181 [nt] 526 424 365 319 282 247 216 189 164 140 1189673523211–7 –26 –46 –68 –89 –110 –131 –150 –168 –185 –202 –222 –245 –266 –283 –305 –327 –352 –377 –398 –409 –436 52°30'52°30'53°30'53°30' 66 °5 0' 66 °5 0' 66 °4 0' 66 °4 0' 66 °3 0' 53°53° a b c d dd e e f f g g h h i sisimiut 10 km qeqerta lik qeqerta lik itilleqitilleq kkaannggeerrlluu aarrssuukk iikkeerrttooooqq mm aalliiggaaaaqq j a b 52°30'53°30' 66 °5 0' 66 °4 0' 66 °3 0' 53° a b c d d d e e f f g g h h isisimiut 10 km qeqerta lik itilleq kangerlu arsuk ikertooq m aligaaq post-dyke (nagssugtoqidian) granulite facies post-dyke (nagssugtoqidian) amphibolite facies pre-dyke (archaean) granulite facies pre-dyke (archaean) amphibolite facies pre-dyke amphibolite facies j fig. 2. correlation between metamorphic facies and aeromagnetic anomaly patterns in the itilleq–ikertooq region. white lines indicate approximate metamorphic facies boundaries based on geological field work; labels a–j are explained in the text. a: distribution and relative ages of metamorphic facies. b: total intensity magnetic field anomaly map. shadow of magnetic field pattern modelled from a light source with inclination 45° and declination 315°. cross-cut gneiss structures. the main dyke direction is nne–ssw, and a subordinate direction is e–w to ese– wnw (fig. 1). upon entering the itilleq area, the dyke trends are e–w, parallel to the fjord. this change in trend also corresponds to a change in foliation trend in the host gneisses. however, the dykes are still largely undeformed and unmetamorphosed within this e–w trend. the metamorphic grade of host gneisses north and south of itilleq is granulite facies in western parts and amphibolite facies in eastern parts (fig. 2a). however, all along the e–w trend in itilleq, gneisses are in low amphibolite facies. the dyke behaviour in the itilleq region led to the interpretation that prior to intrusion of the kangâmiut dykes the area was stabilised in amphibolite-granulite facies with a variable northerly trend of the foliation (grocott 1979; korstgård 1979). at some point prior to dyke intrusion an e–w zone of strong deformation was established along itilleq, downgrading gneisses to low amphibolite facies (epidote-muscovite). within this itivdleq shear zone, dykes intruded along the shear fabrics and show a variety of primary pinch-and-swell structures (nash 1979). outside the shear zone, dyke margins are straight-sided indicating that dykes intruded along brittle fractures. 182 post-dyke metamorphism and deformation farther north of itilleq, from kangerluarssuk and northwards (fig. 2a), dykes are thoroughly deformed and parallel to country rock structures. both dykes and country rock structures are in amphibolite facies. foliation trends are variable ene–wsw around west-plunging fold axes. continuing northwards the metamorphic grade increases and reaches granulite facies north of ikertooq fjord (fig. 2a). in addition, gneiss structures and metamorphosed dykes take on a pervasive e–w orientation (ikertôq thrust zone, fig. 1) with a steeply n-dipping foliation and nplunging stretching lineations. the interpretation of field observations in the northern sno is that the metamorphism and deformation are post-dyke, the metamorphic transition is prograde, and the ikertôq thrust zone represents a zone of southward ductile thrusting whereby deeper-seated rocks are brought up from the north. facies transitions within the itilleq–ikertooq region four types of facies transitions or boundaries are recognised. two of these are prograde and two are associated with strong deformation in ductile shear zones. the amphibolite–granulite facies transition in the archaean areas around itilleq is prograde and static in the sense that the boundary was not established as a result of a deformational event, but reflects static equilibration of the mineral assemblages to the conditions that prevailed when the rocks were at their deepest crustal level. during later uplift the rocks escaped any significant metamorphic changes due to the absence of deformation, and the metamorphism reflects their initial archaean state. the granulite to low amphibolite facies and amphibolite to low amphibolite facies transitions along itilleq are retrograde and dynamic in the sense that they were established as a direct consequence of the deformation along the itivdleq shear zone. mineral assemblages in the shear zone were equilibrated to the metamorphic conditions of a higher crustal level than reflected in the surrounding gneisses, and the shearing triggered this re-equilibration. the amphibolite–granulite facies transition north of ikertooq is both prograde and dynamic. it can be considered as a displaced prograde and static transition brought up into a sub-vertical position by the overthrust movement along the ikertôq thrust zone (fig. 1). magnetisation comparing the magnetic anomaly map for the area (fig. 2b) with the metamorphic map (fig. 2a) a striking coincidence of magnetisation and metamorphic boundaries is evident. more information on the magnetic field data and the geological interpretations can be found in rasmussen & van gool (2000), nielsen (2004) and nielsen & rasmussen (2004). strong magnetisation in pre-dyke archaean granulite facies areas just north of itilleq (a in fig. 2b) is attributed to a higher content of magnetite or other magnetic minerals. a likely explanation for this is production of magnetite by the breakdown of hydrous (fe, mg)-al-silicates (e.g. biotite, amphibole) during the transition from amphibolites to granulite facies according to the general reaction: hydrous (fe, mg)-al-silicates ± sio 2 ± o 2 = kfeldspar + (fe, mg)-silicates ± magnetite + h 2 o. the lower magnetisation in pre-dyke archaean amphibolite facies areas (b in fig. 2b) relative to pre-dyke archaean granulite facies areas indicates no additional production of magnetite. the gradual increase in magnetic intensity (c in fig. 2b) marks the gradual prograde facies transition. the elongate low magnetic anomaly coincident with the itivdleq shear zone (d in fig. 2b) is caused by extensive breakdown of magnetic minerals. this may be due to chemical breakdown during metamorphic retrogression to pre-dyke amphibolite facies aided by circulating fluids in the shear zone, and mechanical destruction of the magnetic mineral grains. the abrupt changes in anomaly patterns from d to a (fig. 2b) across the metamorphic facies transition and deformation boundary are a response to the dynamic nature of this boundary. previously suggested possible shearing south of ikertooq (e in fig. 2b; grocott 1979; korstgård 1979) contemporaneous with the shearing at itilleq (d in fig. 2b) is supported by similarities in the character of the anomaly patterns. the post-dyke amphibolite facies areas at, and south of, ikertooq (f in fig. 2b) indicate the palaeoproterozoic retrogression to amphibolite facies and deformational reworking. the boundary between the pre-dyke archaean amphibolite facies and the post-dyke amphibolite facies areas does not have a well-defined magnetic signature (between b and f in fig. 2b). the increase in magnetisation north of ikertooq (g in fig. 2b) corresponds to rocks metamorphosed under granulite facies conditions after dyke intrusion and brought up by overthrusting. the offset between the mapped facies boundary north of ikertooq (fig. 2a) and the boundary between high and low magnetisation (h in fig. 2b) can be explained as partially due to non-exposed post-dyke 183 granulite facies rocks, and partially to the effect of stacked thrust panels of post-dyke amphibolite and granulite facies rocks with alternating low and high magnetic intensity anomalies (i in fig. 2b). isolated high intensity anomalies can be correlated with distinct lithologies or intrusives (e.g. an anorthosite complex at j in fig. 2b). the presence or absence of kangâmiut dykes is not reflected in the aeromagnetic data. the observed correlations between metamorphic facies, deformation and magnetisation can be extended to other areas of the sno (fig. 3) provided that the background gneisses are lithologically fairly homogeneous, as is generally the case in the southern nagssugtoqidian orogen. where gneiss lithologies are more variable, such as in the nordre isortoq steep belt (fig. 1) and the nordre strømfjord shear zone (sørensen et al. 2006, this volume) correlations tend to depend on lithology rather than metamorphic grade. acknowledgements the authors thank graham leslie and chris pulvertaft for their concise and constructive reviews. references escher, j.c. & pulvertaft, t.c.r. 1995: geological map of greenland, 1:2 500 000. copenhagen: geological survey of greenland. grocott, j. 1979: controls of metamorphic grade in shear belts. in: korstgård, j.a. (ed.): nagssugtoqidian geology. rapport grønlands geologiske undersøgelse 89, 47–62. korstgård, j.a. (ed.) 1979: nagssugtoqidian geology. rapport grønlands geologiske undersøgelse 89, 146 pp. marker, m., mengel, f., van gool, j. & field party 1995: evolution of the palaeoproterozoic nagssugtoqidian orogen: dlc investigations in west greenland. rapport grønlands geologiske undersøgelse 165, 100–105. nash, d. 1979: an interpretation of irregular dyke forms in the itivdleq shear zone, west greenland. in: korstgård, j.a. (ed.): nagssugtoqidian geology. rapport grønlands geologiske undersøgelse 89, 77–83. nielsen, b.m. 2004: crustal architecture and spatial distribution of mineral occurrences in the precambrian shield of central west greenland based on geophysical and geological data. danmarks og grønlands geologiske undersøgelse rapport 2004/26, 63 pp., 8 appendices. ph.d. thesis 2004. department of earth sciences, university of aarhus, denmark. nielsen, b.m. & rasmussen, t.m. 2004: mineral resources of the precambrian shield of central west greenland (66° to 70°15′n). part fig. 3. total intensity magnetic field anomaly map of the south-eastern part of the nagssugtoqidian orogen and its foreland, with the location of the itilleq–ikertooq region (white frame, fig. 2). abbreviations as for fig. 1; shadow on magnetic data as for fig. 2. the e–w-trending thrust zone with question marks north of kangerlussuaq is uncertain, as this structure has not been confirmed by geological mapping. sn f southern nagssugtoqidian foreland (north atlantic craton) n ag ss ug to qi di an o ro ge n nordre iso rtoq steep belt southern cno ikertôq thrust z one itivd leq shear z one c n o sn o 50°50°52°52° 66 °4 0' 66 °4 0' 67 °0 5' 67 °0 5' 66 °1 5' 66 °1 5' sø nd re str øm fjo rd sø nd re str øm fjo rd sisimiutsisimiut kangerlussuaqkangerlussuaq itilleqitilleq ikertooq ikertooq qeqertalikqeqertalik kangerluarsukkangerluarsuk fig. 2 53°53° [nt] ?? 50 km 536 407 334 284 243 209 178 151 125 76 54 32 11 –9 –30 –50 –72 –91 –110 –127 –145 –161 –176 –190 –204 –219 –234 –264 –282 –303 –326 –354 –386 –416 –461 –535 184 3. implications of potential field data for the tectonic framework. danmarks og grønlands geologiske undersøgelse rapport 2004/ 21, 165 pp. nutman, a.p., kalsbeek, f., marker, m., van gool, j.a.m. & bridgwater, d. 1999: u-pb zircon ages of kangâmiut dykes and detrital zircons in metasediments in the palaeoproterozoic nagssugtoqidian orogen (west greenland): clues to the pre-collisional history of the orogen. precambrian research 93, 87–104. ramberg, h. 1949: on the petrogenesis of the gneiss complexes between sukkertoppen and christianshaab, west greenland. meddelelser fra dansk geologisk forening 11, 312–327. _____________________________________________________________________________________________________________________________________________________________ manuscript received 10 november 2004; revision accepted 1 november 2005 rasmussen, t.m. & van gool, j.a.m. 2000: aeromagnetic survey in southern west greenland: project aeromag 1999. geology of greenland survey bulletin 186, 73–77. sørensen, k., korstgård, j.a., glassley, w.e. & stensgaard, b.m. 2006: the nordre strømfjord shear zone and the arfersiorfik quartz diorite in the inner arfersiorfik, the nagssugtoqidian orogen, west greenland. in: garde, a.a. & kalsbeek, f. (eds): precambrian crustal evolution and cretaceous–palaeogene faulting in west greenland. geological survey of denmark and greenland bulletin 11, 145–161 (this volume). geological survey of denmark and greenland bulletin 26, 2012, 49-52 49© 2012 geus. geological survey of denmark and greenland bulletin 26, 49–52. open access: www.geus.dk/publications/bull groundwater protection in denmark and the role of water supply companies jacob dyrby petersen and lisbeth flindt jørgensen denmark has a decentralised water supply structure with about 2500 water supply companies. until recently, about 150 of these, especially the larger ones, were owned by local authorities; the rest are private, all run on an independent and not-for-profit basis. recently, a new law, the water sector law (miljøministeriet 2009), was implemented. its purpose is to privatise the water supply sector (although, as hitherto, into not-for-profit corporations), and statutory duties are separated from operations in order to make the supply of drinking water to consumers as efficient as possible. an important element of the water sector law is the introduction of a new regulatory body, the utility secretariat. the role of this new institution under the danish competition and consumer authority is to enforce price ceilings on drinking water, based on a selection of benchmark parameters. danish policy is to base the drinking water supply on unpolluted groundwater. nevertheless, in 2010, pesticides were detected in 25% of all tested, active abstraction wells, and the level for individual substances was exceeded in 4.5% of the wells (thorling et al. 2011), a situation more or less unchanged over the last 20 years. this calls for further and continuous measures to strengthen groundwater protection. in addition, the range of protective activities implemented over the last decades may require reinforcement as groundwater monitoring data show that a high proportion of young groundwater is polluted with pesticides (thorling et al. 2011). our hypothesis is that during the last decades groundwater protection has changed from being a national responsibility to becoming a more locally embedded task. we also see a change towards the water supply companies being the major actors capable of initiating active groundwater protection. this transition also implies a change as regards economy, viz. from public authorities to consumers. the aim of this paper is to discuss how the water sector law affects the frames for groundwater protection when an increasing part of the effort hinges on the water supply initiative taken by private companies. background denmark is one of the countries in the world most heavily reliant on groundwater as more than 99% of all water supplies are derived from this source (jørgensen & stockmarr 2009). thus denmark has a strong incentive to, and tradition for, knowledge-based groundwater management. the characteristics of the danish hydrogeological conditions, combined with wide-ranging groundwater monitoring and a relatively advanced technological and administrative approach as well as a tradition for public involvement in administration, have shaped the present framework of groundwater protection in denmark. in denmark, official groundwater protection began in the 1970s with the creation of the ministry of the environment. the ensuing environmental legislation and the establishment of 14 regional authorities prepared the path for regional groundwater management strategies based on national legislation for the protection of the environment, including groundwater. as a consequence of a statutory framework reform in 2007, abolishing the 14 regional authorities, groundwater protection was transferred to the 98 municipalities (kommuner). initially, mainly prospective activities were addressed, but a few years later remediation acts were enforced and resources allocated to manage the ‘sins of the past’ in terms of old polluted sites. water supply and water demand manage-a modern danish waterwork from østerby, eastern jylland. 5050 ment strategies emerged, e.g. by launching saving campaigns targeted at both industrial and private consumers. appeals soon followed not to use pesticides in private gardens etc. to reduce groundwater contamination. these public campaigns were initiated and implemented by the state as well as the water supply companies. the latter were allowed a more operational role in groundwater protection by a change of the water act in 1999, financing activities through water levies (miljøministeriet 1999). thus groundwater protection was no longer solely dependent on initiatives by the state. methods to analyse the past and present status of involvement in groundwater protection and to examine what is conceived as the overriding issues of the new water sector law, two major sources are used: (1) a survey conducted in 2010 among the 75 largest danish water supply companies (pedersen 2010a) and (2) qualitative interviews in 2011 with key persons in the water sector (table 1), combined with a concurrent study of the political and legislative process surrounding the implications of the water sector reform (petersen 2011). given the dynamic nature of the field of study, the interviews were conducted within a period of six months with follow-ups carried out as semi-structured interviews based on an iterative and adaptive approach (kvale 1997). the interviews covered the views of informants on: • the historical development of the water supply sector; • the interplay between the authorities and the water supply companies; • the eff orts and players in the danish groundwater protection regime and • the legislative process and the implications for the economic and ecological aspects of future groundwater protection. the theoretical frame was partly based on a theoretical network analysis adapted from sørensen & torfing (2005) and partly on a meta-analysis of the understanding of groundwater management from the definition of market environmentalism (bakker 2005). the network approach was used to analyse the use of decision making and coordinating networks between actors instead of the two traditional management instruments: hierarchy (i.e. state control) and free market. this enables an assessment of how the water sector law has changed the opportunities of the water supply sector to form and co-operate within networks. market environmentalism can be used as a theoretical tool to analyse the development of resource understanding and resource management based on market mechanisms, based on studies from the uk. from being a supply-led, state-owned and managed resource, groundwater in the uk is increasingly governed through the private sector and market-based instruments. as a country of comparison, the theoretical framework is used to analyse the development in denmark. results in 2010, 48% of the 75 largest danish water supply companies were actively involved in groundwater protection activities, and were partly financing these (pedersen 2010a). the companies reported that they had used between 1 and 25% of their annual turnover on groundwater protection (pedersen 2010b). although these estimates are highly uncertain, water supply companies have reported increasing expenditure on these protection activities (fig. 1), and the figures indicate interviewed institution representing two large, urban water supply companies water supply two water supply associations professional body one agricultural association professional body two municipalities (kommuner) public sector (local) the nature agency (ministry of the environment) public sector (national) the utility secretariat (danish competition and consumer authority)* public sector one large, private contractor/consultant private sector table 1. the interviewed institutions * the authority implementing the water sector law 0 40 80 year 2000 e x p e n d it u re (m ill io n d k k ) 120 2005 2010 fig. 1. expenditure related to groundwater protective activities as reported by the water supply companies. modified from petersen (2011). 51 a total annual funding from the water supply sector of more than 100 million danish kroner (petersen 2011). more than half of the interviewed water supply companies expected expenditure to increase or at least remain at the same level in forthcoming years, and more than 60% expected necessary future activities (pedersen 2010b). the interviews conducted by petersen (2011) reflect that the water supply companies have taken care of a range of different tasks connected with groundwater protection. they have often voluntarily entered networks (see fig. 2) with each other and the relevant municipalities and thereby broadened the character of tasks to a degree that would otherwise have been impossible, based on individual water supply company economy and human resources. the tasks cover afforestation, volunteer farming agreements with landowners, groundwater resource and quality investigations, entry into action plans together with other water suppliers, stakeholders and local authorities, etc. these tasks were previously financed by adding an extra consumer levy on the water price, typically a groundwater protection tariff of less than 2% of the total water price per cubic metre. as a consequence of the water sector law, the water supply companies fear that this will not be possible in the future as the law allows the utility secretariat to establish differentiated price ceilings for the individual water supply companies (petersen 2011) based solely on production expenditure. one of the two water supply organisations have expressed their concerns at this new development, expecting that the water supply companies will have to focus their activities only on economic issues and not, as previously, be able to perform groundwater protective activities for long-term consumer benefit (petersen 2011). the interviews revealed concerns regarding the capability of local authorities to set up groundwater protective activities at the same level as the former regional authorities, due to limited resources under the new statutory framework. these concerns are related to both spatial (arm’s length) aspects (the water resource being managed at a local and not a regional level, giving local interests larger influence) and resource issues (the local authorities lacking possibilities of sustaining both economy and skills; petersen 2011). furthermore, there is fear among water associations that the local authorities will only have an active role in protecting the groundwater in quantitative terms, as qualitative aspects are now to be handled in national legislation and action plans, and not at a local level. thus the municipalities may only be in a position to administer and outline the overall framework, and only perform limited actual groundwater protective activities. accordingly, the interviewed municipalities expressed their expectations that the water supply companies are to be the active players in these activities, and that the economic resources are to be found through water levies (petersen 2011). the capacities of the water supply companies should therefore include the practical implementation of future groundwater protection measures, at least as seen from the municipalities’ point of view. discussion over the past two decades, the water supply companies have entered into different networks.these networks have been established for various reasons, which could be an experience of missing coordination in societal management, or a need for mutual action and sharing of knowledge and resources (sørensen & torfing 2005). the networks have been associated especially with groundwater protective activities such groundwater protection groundwater protection water supply company b water supply company d water supply company c water supply company a utility secretariat municipality municipality water supply company after water sector law before water sector law fig. 2. the framework for groundwater protection before and after the implementation of the water sector law. the networking is expected to decrease, and the water supply companies will change from being both operators, planners and decision makers to be operators only. 5252 as afforestation, voluntary agreements with landowners on non-pesticide cultivation, investigations and action plans (petersen 2011). how does the new water sector law influence these networks and their associated activities? the law sets strict limits on how the water supply companies can use their financial resources. the first few trials have already shown that the water sector law does not provide the framework necessary for continuation of former activities such as afforestation and networking, as the financial platform in terms of a water tax is not in accordance with the law. both the water supply companies and the local authorities look upon this as a bottleneck in establishing and continuing effective groundwater protection (petersen 2011). the intension behind the water sector law is to ensure economically efficient company operations and at the same time the law emphasises that the main objective of the companies is to abstract, treat and distribute drinking water (petersen 2011). on the other hand, a central element is that the water supply companies are allowed – if not obliged – to perform groundwater protective activities if required by the local authorities as part of a public environmental action plan, dictated by the central authorities. as the interviews revealed, the experience and financial capacity of the decentralised water supply companies have played an essential role in the protection of the danish groundwater over the past decades, and there is considerable concern regarding the future fate of these activities (petersen 2011). the new regulatory paradigm is by bakker (2005) described as market environmentalism, where market mechanisms are to ensure efficiency, competition and sustainability. whereas the business aspects are beyond the scope of this paper, the environmental perspective implies a movement towards a capitalisation of the environment to include resource protection in water pricing, as attempted in the uk. from a quantitative perspective, the price can reflect most of the cost related to abstraction and distribution. qualitative groundwater protection, however, has entrenched complexities. due to the intertemporal aspect and imperfect, although ever increasing knowledge, the financing of qualitative protection may shift during the coming years as the water supply companies are encouraged to focus more narrowly on initiatives directly related to their resource. our study shows that combined with the recent structural reforms of the environmental management authorities, the water sector law sets the stage for a shift in a formerly accepted managerial practice. this is mainly seen as a drawback as it reduces the possibilities of the water supply companies to undertake groundwater protective actions that are seen as necessary by the individual companies and leaves them in a role where they are only operators of actions dictated by the municipality. the benefits may, on the other hand, be a better coordinated and integrated water management policy, where the groundwater protection is seen in a more holistic approach that also calls for actions in areas where the water supply companies have not so far been active. however, this calls for a strong incorporated and active integrating governance strategy from the local authorities. only the future can show how this will come out. references bakker, k. 2005: neoliberalizing nature? market environmentalism in water supply in england and wales. annals of the association of american geographers 95, 542–565. jørgensen, l.f. & stockmarr, j. 2009: groundwater monitoring in denmark: characteristics, perspectives and comparison with other countries. hydrogeology journal 17, 827–842. kvale, s. 1997: interview. en introduktion til det kvalitative forskningsinterview, 318 pp. københavn: hans reitzels forlag. miljøministeriet 1999: lovbekendtgørelsen nr 130 af 26. februar 1999 om vandforsyning mv. miljøministeriet 2009: lovbekendtgørelsen nr 469 af 12. juni 2009 om vandsektorens organisering og økonomiske forhold. pedersen, p.e. 2010a: vandforsyninger under hårdt pres. vækst 131(2), 12–13. pedersen, p.e. 2010b: et væld af metoder til grundvandsbeskyttelse. vækst 131(2). only available online: http://www.hedeselskabet.dk/page.6360. aspx?recordid6360=335 petersen, j.d. 2011: dyrebare dråber. vandforsyningernes rolle i grundvandsbeskyttelse, 86 pp. unpublished master thesis, roskilde universitet, danmark. sørensen, e. & torfing, j. 2005: netværksstyring: fra government til governance, 218 pp. frederiksberg: roskilde universitetsforlag. thorling, l., hansen, b., langtofte, c., brüsch, w., møller, r.r., mielby, s. & højberg, a.l. 2011: grundvand. status og udvikling 1989–2010, 139 pp. særudgivelse. københavn: de nationale geologiske undersøgelser for danmark og grønland. only available online: http://www. geus.dk/publications/grundvandsovervaagning/1989_2010.htm authors’ addresses j.d.p., roskilde university, department of technological and socio-economic planning, universitetsvej 1, dk-4000 roskilde, denmark. e-mail: dyrby@ruc.dk l.f.j., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. geological survey of denmark and greenland bulletin 15, 2008, 41-44 in 2006, dong energy initiated the development of the horns rev ii offshore wind farm in the north sea (fig. 1). in order to evaluate and map the characteristics of the surface features of the sea bed and to characterise the subsurface in the wind farm area, the geological survey of denmark and greenland (geus) conducted a geophysical survey of the area. the survey utilised a variety of instruments: sparker, side-scan sonar, marine caesium magnetometer and a multibeam echo-sounder. in addition, information on the subsurface sediments was obtained by cone penetration tests (cpt) and by drilling to 30–50 m below the sea bottom. geological correlation of the cpt results with the other survey results was extremely complicated but was required in order to understand the architecture of the ice marginal glaciotectonic complex. information on the geology is crucial for evaluation of the geotechnical problems of the region. methods shallow seismic, multibeam and coring methods are well established methods in marine geological research. in contrast, marine magnetometer and cpt studies are normally used to find iron objects and for the determination of geo technical properties. however, both magnetometer and cpt data can yield valuable geological information, as de scribed below. the magnetometer, which gives information about the ambient magnetic field strength by measuring the variation in caesium electron energy level, is towed behind the survey vessel. anomalies in the magnetic field are related to lithological differences in the area. the gridded variations in the strength of the magnetic field, as measured along the survey lines, re vealed a spatial pattern. some areas show homogeneous field strength, whereas other areas show a heterogeneous pattern of field strength. initially, this pattern appeared to be spatially complex. however, we found that there is a close connection between the magnetic data and large-scale glaciotectonic deformations, as shown by comparisons with other geological and seismic data. the cpt is carried out using a cylindrical penetrometer with a conical tip (cone) penetrating the ground at a constant rate (2 cm s–1). during the penetration, the forces on the cone and the friction sleeve are measured. the cpt results at horns rev aid in the foundation design for each of the 98 windmills. general geology horns rev can be divided into an inner and outer part separated by a 20 m deep channel, slugen (fig. 1). a reef con© geus, 2008. geological survey of denmark and greenland bulletin 15, 41–44. available at: www.geus.dk/publications/bull 41 10 15 15 15 15 15 10 10 10 25 15 20 25 10 20 20 20 10 20 20 20 20 28 10 10 10 10 10 10 15 10 10 4 6 6 6 6 6 6 4 4 4 6 4 esbjerg 10 km vejers bank slugen blåvands huk outer horns rev inner horns rev fanø 1010 15 15 10 25 25 10 20 20 20 20 10 10 10 15 6 6 6 4 6 15 15 15 20 20 vejers strand fanø 7°30′e 55°40′n 55°50′n 55°30′n 8°00′e 8°30′e hill island lake/bog deposits marine deposits aeolian deposits limit of survey area cone penetration test vovov denmark north sea 125 km fig. 1. map of the horns rev region showing the survey area north of the outer part of horns rev. the plan is to build 98 windmills. the survey area covers approximately 58 km2 with a length of 12 km from south to north and a width of 6 km from east to west. modified from larsen (2003). geology of outer horns rev, danish north sea jørn bo jensen, peter gravesen and steen lomholt sisting of holocene shallow submarine sandbanks exists at a water depth of 2–7 m, whereas glacial and interglacial de posits occur in the surrounding area at a water depth of 20–30 m. tertiary strata, most likely of miocene age, comprise the basement to the quaternary deposits located >50 m below the sea bed. the wind farm prospect area is located north of the outer horns rev (fig. 1), at a water depth of 7–18 m on the vovov hill island (bakkeø). previous investigations have shown that the horns rev area was ice-covered during the elsterian and saalian glaciations (larsen & andersen 2005). marine late elsterian – holsteinian sediments that were deposited before the saalian have also been identified in the central north sea (long et al. 1988) as well as onshore south-western jylland (knudsen 1987, 1994). eemian marine interglacial sediments have also been recorded in the region (konradi et al. 2005). saalian glacial deposits and other older quaternary deposits the detailed seismic survey results coupled with coring and cpt results (fig. 2) show that it is possible to identify elsterian glacial clay and sand till deposits that are overlain by holsteinian interglacial marine clay, as confirmed by fora miniferal analyses. the top of the glacial deposits is characterised by a regional erosion surface that appears as a notable seismic unconformity (figs 2, 3). this reflector represents the surface of the vovov hill island. below the unconformity, chaotic seismic reflectors record the existence of small-scale hummocky clinoforms, vertical and sub-horizontal reflectors and channel-like features (fig. 2). similar structures that have been recognised in the southern north sea have been interpreted as representing the ice margin of the late saalian warthe advance (andersen 2004). on vovov hill island, the glacial top surface generally lies 18–20 m below sea level (b.s.l.). however, at the north-eastern rim of the wind farm where the western margin of the horns rev valley is situated, the surface drops to 39 m b.s.l. the saalian glacial deposits consist mainly of mediumand coarse-grained sand and gravel with subordinate finegrained sand and silt layers. in addition, fine-grained and silty sand with mica and plant fragments also occurs. this composition is similar to that of the onshore hill islands in western jylland, where the deposits are also very sandy (sjørring 1981). sandy and clayey tills and diamictons with gravels, stones and boulders have also been found. the chaotic seismic image suggests the presence of glaciotectonically deformed sand and gravel (and tills) that were dislocated at the end of the saalian glacial stage (fig. 2). the seismic sections also show several channels and valleys that have incised into the glacial hill (fig. 3). the reflectors inside these depressions have parallel features that suggest the presence of water-lain sediments. the valleys, therefore, were filled with glaciofluvial sand and gravel, probably during the last phase of the late saalian warthe advance and during decay of the ice. it is important to note, however, that the top sediment deposits in these valleys may also be younger in age. eemian deposits earlier studies of foraminiferal faunas from sediment cores suggested that eemian deposits occur in the horns rev area. however, the seismic evidence for eemian sediments is based on generally indistinct seismic reflectors that are often capped by a sharp top reflector (konradi et al. 2005). the eemian sediments are up to 13 m thick, with the top of the unit located 11–14 m b.s.l. the deposition of the eemian unit corresponds to a sea-level high-stand during the eemian period when the sea covered almost the entire area. the only units not entirely transgressed by the sea were the saalian glacial deposits at the vovov hill island. in the wind farm area, marine eemian deposits have been identified using biostratigraphical data. the deposits are olive-grey silty clay and sandy silt with sand lenses. the sediments are often bioturbated and contain shells and shell fragments. the marine eemian layers form a wedge that laps onto the vovov hill island. patchy channel infill occurs in the central part of the wind farm area (fig. 3). eemian freshwater lake deposits are found below the marine layers. 42 5 m 60 50 40 tw ow ay t ra ve l t im e (m se c) 30 20 10 borehole cpt 5 50 m sea bed holocene marine sand multiplemultiplemultiple holsteinian marine clay saalian glacial sand elsterian till fig. 2. seismic example from sparker line 118 showing correlation to borehole and cpt data. holocene deposits due to the absence of weichselian de posits, holocene marine deposits directly overlie the vovov hill island regional erosional surface (fig. 3), as illustrated in the seismic pattern by internal horizontal layering (fig. 2). the surface cuts into the saalian de posits and also into the eemian marine deposits when present. the holocene marine deposits form a relatively thin sand cover over all of the glacial and interglacial sediments. in small depressions, early holocene freshwater se diments with plant remains are occasionally noted. the thick ness of the holocene sand layers ranges from 1 to 2 m in most parts of the vovov hill island. in contrast, the thickness of the holocene deposits increa ses to 6 to 8 m in the northern margin of the wind farm area. hill island glacial deformations detailed seismic mapping reveals that the wind farm area is located in a complicated ice margin zone of late saalian age. the ridges were formed by ice push from the east (fig. 4). con sequently, the wind farm area is characterised by ice marginal deformations that exhibit correlation with the gridded variations in magnetic field strength along the survey lines (fig. 4). saalian glacio-fluvial sediments were deposited to the west of the glacier margin and are only weakly disturbed by glacial ice push. the deformations associated with the initial ice push mainly affected the central part of the wind farm area (figs 3, 4). the western parts of the survey area are characterised by a veneer of holocene deposits, which are underlain by non-disturbed deposits. on the magnetometer map the western survey areas show no magnetic anomalies due to the simple geology (fig. 4). the eastern two thirds of the wind farm area are influenced by glacial deformation in the form of thrusting and folding. in the folded areas, the seismic reflectors have a char acteristic wavy appearance (figs 2, 3). it is in these areas that holsteinian, interglacial clay deposits have been found folded into positions close to the sea bed in borehole 5 (fig. 2). it is believed that soft miocene clay acted as a décollement layer. on the magnetometer map, the lack of strong magnetic anomalies makes it possible to distinguish folded areas from severely glacially deformed areas (fig. 4). severely disturbed areas interfinger with the folded areas, in a general sw–ne direction (fig. 4). on the seismic profiles, the reflections are typically chaotic, with the deeper re flectors disappearing completely (fig. 3). consequently, the sediment layering would be expected to be chaotic, with many changes in lithology. the magnetometer map shows that anomalies are common in the severely disturbed areas, yielding an uneven map surface that is clearly related to the chaotic lithological distribution (fig. 4). elongated depressions with a nw–se trend are filled with fluvial deposits (fig. 4). examination of these depressions and their fill, as demonstrated in seismic line 102 (fig. 3), shows that both the fluvial deposits and the channels follow the synclinal depressions of the folded areas. on the magnetometer map (fig. 4), the fluvial deposits can be identified by relatively large nw–se-trending anomalies. glacial deformation style the combination of deformation classes reveals a sw–netrending border, separating deformed and undeformed deposits (figs 3, 4). this distribution reveals an ice-push direction from the east and the holsteinian marine sediments involved suggest that the deformations are associated with the late saalian warthe ice margin, as no younger ice ad vance reached the area. the horns rev ice-margin push moraines can be classified as thin-skinned fold and thrust deformations following bennett (2001). the large and welldeveloped fold deformation structures indicate that relatively high glacial stress (gravity spreading and/or push) was in volved and that miocene clay layers acted as an important décollement horizon. outside the ice margin, the rather high foreland strength of the warthe outwash fan resulted in a marginal zone dominated by thrust deformations. to the west only weakly deformed outwash fan deposits are found. 43 500 m 10 50 500 m eastwest cpt 44 massive glaciotectonised zone glacial thrust zone holocene marine ? ? ? ? ? ? ? 10 d ep th ( m ) d ep th ( m ) 50 deformation classes eastwest a b cpt 44 glacial thrust zone? ? ? ? ? ? ? eemian marine/saalian fluvial saalian glacial elsterian glacial holsteinian marine internal reflector pre-quaternary reflector? weakly glacial-disturbed deposits glacial-folded deposits heavily glacial-disturbed deposits channel deposits internal reflector pre-quaternary reflector? fig. 3. interpreted sparker line 102 illustrating (a) the seismic stratigraphy and (b) the different types of glaciotectonic deformation classes. for location see fig. 4. 44 concluding remarks sediment cores and cpt tests from the horns rev ii wind farm area show that the region is characterised by a complex geology. large-scale glaciotectonic deformations are related to the late saalian warthe advance with ice push from the east. this investigation also shows that there is a close correlation between magnetometer data and seismically mapped types of deformation structures. the glacial push moraines developed under conditions of high glacial stress, with miocene clay acting as décollement horizon. it is important to note that the new evidence from outer horns rev moves the position of the warthe ice margin more than 50 km further to the west than formerly suggested. acknowledgement we thank dong energy for the opportunity to present the geological results from the horns rev ii offshore wind farm project. references andersen, l.t. 2004: the fanø bugt glaciotectonic thrust fault complex, southeastern danish north sea. ph.d. thesis 2004. danmarks og grøn lands geologiske undersøgelse rapport 2004/30, 143 pp. bennett, m.r. 2001: the morphology, structural evolution and significance of push moraines. earth science reviews 53, 197–236. knudsen, k.l. 1987: foraminifera in the late elsterian – holsteinian sequence at tornskov in south jutland, denmark. danmarks geolo giske undersøgelse serie b 10, 7–31. knudsen, k.l. 1994: the marine quaternary in denmark: a review of new evidence from glacial – interglacial studies. bulletin of the geological society of denmark 41, 203–218. konradi, p.b., larsen, b. & sørensen, aa.b. 2005: marine eemian in the danish eastern north sea. quaternary international 133, 21–31. larsen, b. 2003: blåvands huk – horns rev området – et nyt skagen? geologi – nyt fra geus 4, 12 pp. københavn: danmarks og grønlands geologiske undersøgelse. larsen, b. & andersen, l.t. 2005: late quaternary stratigraphy and morphogenesis in the danish eastern north sea and its relation to onshore geology. geologie en mijnbouw 84, 113–128. long, d., laban, c., streif, h., cameron, t.d.j. & schüttenhelm, r.t.e. 1988: the sedimentary record of climatic variation in the southern north sea. philosophical transactions of the royal society of london b318, 523–537. sjørring, s. 1981: pre-weichselian till stratigraphy in western jutland, denmark. mededelingen rijks geologische dienst 34, 62–68. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: jbj@geus.dk limit of survey area fluvial deposits heavily glacial-disturbed glacial-folded weakly glacial-disturbed line 102 line 118 5 2.5 km 49710 nt 49740 nt 49760 nt 49780 nt 49800 nt 49820 nt 49860 nt borehole 2.5 km fig. 4. map distribution of the different types of glaciotectonic deformation structures and magnetometer anomalies. the locations of sparker line 102 and borehole 5 are shown. geological survey of denmark and greenland. bulletin 10, 49-52 until recently, in situ u-pb zircon geochronology could be carried out only using ion microprobes, requiring lengthy analysis times of c. 20 minutes. however, new developments in laser ablation inductively coupled plasma mass spectrometer technologies have resulted in zircon geochronology techniques that are much faster, simpler, cheaper, and more precise than before (e.g. frei et al. 2006, this volume). analyses approaching the precision obtained via ion microprobe can now be undertaken in 2–4 minutes using instruments such as the 213 nm laser ablation (la) system coupled with element2 sector-field inductively coupled plasma mass spectrometer (sf-icp-ms) housed at the geological survey of denmark and greenland (geus). the up to tenfold decrease in analytical time means that zircon geochronology can now be used in a much wider range of studies. the godthåbsfjord region, southern west greenland, contains some of the oldest rocks exposed on the earth’s surface reflecting a very complex archaean geological evolution (figs 1, 2). over recent years geus has undertaken a range of mapping projects at various scales within the godthåbsfjord region (see also below). these include the mapping of the 1:100 000 scale kapisillit geological map sheet (fig. 1), and regional and local investigations of the environments of formation and geological evolution of supracrustal belts, hosting potentially economic mineral occurrences. zircon geochronology is an important tool for investigating a range of geological problems in this region. by breaking down the complex geology into a series of simple problems that can be addressed using this tool, the geological evolution can be unlocked in a stepwise manner. three examples are presented below: (1) the mapping of regional structures; (2) characterising and correlating supracrustal belts; and (3) dating metamorphism and mineralisation. although focus is on the application of zircon geochronology to these problems, it is important to note that the resulting data must always be viewed within a wider context incorporating geological mapping and structural, geochemical and petrographic investigations. regional geology the geology of the godthåbsfjord region is dominated by orthogneiss formed during several distinct episodes of crustal growth during the archaean (fig. 2). these different-aged gneisses are thought to represent distinct small continental blocks that were amalgamated during the neoarchaean (at © geus, 2006. geological survey of denmark and greenland bulletin 10, 49–52. available at: www.geus.dk/publications/bull using zircon geochronology to resolve the archaean geology of southern west greenland julie a. hollis, dirk frei, jeroen a.m. van gool,adam a. garde and mac persson 49 fig.3 c. 3.87-3.6 ga c. 3.0-2.98 ga c. 3.2-2.98 ga c. 3.87-3.6 ga & c. 2.835-2.75 ga c. 2.92-2.84 ga greenland supracrustal belts terrane boundary 25 km nuuk st or ø kapisillit 64°n 51°w52°w 50°w 64°30' n godthåbsfjord ameralla ameralik fig. 1. overview map of the godthåbsfjord region. inset shows the location of the main map in greenland. supracrustal belts are shown in green. the boundaries of the 1:100 000 scale kapisillit geological map sheet area are in red. blue lines outline the area shown in fig. 3. bold black lines are inferred terrane boundaries, with the major age components of the different terranes indicated (after friend & nutman 2005). c. 2.7–2.6 ga) during collisional tectonism (similar to that seen in modern mountain belts). within, and often between, these different crustal blocks, or terranes, supracrustal belts made up of metasedimentary and metavolcanic rocks occur. some of these are known to host potentially economic mineral occurrences, e.g. gold-mineralised supracrustal rocks on the island storø (fig. 1). several high-grade metamorphic events and associated deformation have affected different parts of the region from the palaeoarchaean through to the neoarchaean. these events resulted in partial melting, variable development of high-strain structural fabrics, and several generations of largescale folds. mapping major regional structures although the orthogneisses that dominate the geology of the region were formed during several distinct events, different generations of orthogneisses are often very difficult to distinguish in the field owing to their similarity of composition (fig. 3). the geometry of supracrustal belts is very useful in identifying the nature of regional structures, as these are lithologically very different from the orthogneisses and can be used as structural markers. however, where these are absent reconnaissance geochronology can be applied. having identified several of the main crust-forming events in the region, it is possible to use the emplacement ages of orthogneisses as a guide to field mapping and the delineation of large-scale structures. selected samples collected during the 2004 field mapping were dated using la-sf-icp-ms u-pb zircon geochronology (fig. 2). major fold structures inferred from field mapping data were dated using this method, in some cases resulting in significant advances in understanding. so far the results indicate that mesoarchaean rocks were thrust northwards over neoand eoto palaeoarchaean rocks and then deformed in kilometre-scale refolded folds (fig. 3). the resulting map was used to identify problematic areas to investigate during the 2005 mapping. identifying and correlating supracrustal belts figure 1 shows the distribution of known supracrustal belts in the godthåbsfjord region, typically comprising high-grade metamorphic volcanic, ultramafic, subordinate aluminous and siliceous sedimentary rocks. these were once thought to represent a single, dismembered suite of mesoarchaean supracrustal rocks. it is now known that there are in fact several distinct belts deposited at c. 3.87 and 3.7 ga (the isua supracrustal belt), c. 3.0 ga, and at c. 2.8 ga (fig. 2). these belts commonly occur along tectonised boundaries between chronologically distinct terranes. correlation between different supracrustal belts is important for understanding the regional structures and identifying areas of potentially economic mineral occurrences. some belts show lithological and geochemical differences, and similarities that can be used to compare and contrast them. these indicate that there are at least three different environments that were important in forming the different belts: (1) basic volcanism in extensional oceanic environments; (2) gabbro-anorthosite magmatism at deeper levels within oceanic crust; and (3) andesitic volcanism and associated hydrothermal syngenetic alteration in island-arc environments (garde 2005; hollis 2005). however, the majority of 3.23.6 2.54.0 age (ga) 2.8 eoarchaean palaeoarchaean mesoarchaean neoarchaean ? supracrustals infracrustals metamorphism fig. 2. simplified summary of the major thermal events in the godthåbsfjord region. 1 2 2 1 eoto palaeoarchaean (3.7-3.5 ga) early mesoarchaean (3.2-2.9 ga) late mesoarchaean (2.9-2.85 ga) neoarchaean (2.85-2.75 ga) unknown age colour-coded dated samples first generation folds second generation folds kangersuneq kapisillit ameralla 10 km fig. 3. simplified version of map of major fold structures in the kapisillit 1:100 000 scale geological map area. coloured circles indicate the sites where zircons from rock samples have been dated. the fill colours indicate the obtained ages. these ages have been used to constrain the surface geometry of the folds. photographs of outcrops of orthogneiss from two localities show that rocks of very similar appearance can be very different ages and therefore geochronology is useful in distinguishing them. 50 the belts are strongly tectonised and extensive recrystallisation occurred during high-grade metamorphism. locally hydrothermal alteration is very intense. therefore, the protoliths are difficult to recognise and the different supracrustal belts are difficult to distinguish without detailed geochemical and geochronological information. one of the most important tools for correlation is the study of detrital zircons. siliceous and aluminous metasedimentary rocks carry detrital zircons derived from the erosion of the source regions for the sediments. thus studies of their detrital zircons allow an assessment of both likely sources and their maximum age of deposition (given by the youngest detrital zircon). furthermore, the detrital zircon age spectrum of a metasedimentary rock from a specific supracrustal belt is often characteristic, as the zircons are typically derived from the same source area. the detrital spectra can therefore be used to correlate between widely spaced supracrustal belts that would otherwise be difficult to compare from petrographic or geochemical data alone. a compilation of existing and recently obtained detrital zircon data for metasedimentary rocks and primary zircons from metavolcanic or volcanoclastic rocks in the godthåbsfjord region is shown in fig. 4 (see also hollis 2005). distinct sources for supracrustal rocks of different ages are readily apparent. the volcanic rocks that form part of the mesoarchaean supracrustal belts were deposited at c. 3.07 ga, which is just slightly older than the dominant mesoarchaean regional orthogneiss. similarly, metasedimentary rocks within the mesoarchaean supracrustal belts show dominant zircon age peaks at c. 3.07 ga. however, significant proportions of older zircons show that there was also a contribution from palaeoand eoarchaean sources. the neoarchaean volcanic rocks were deposited at c. 2.83 ga, which is also an age that is well represented in the detrital zircon population of most of the neoarchaean metasedimentary rocks. a few metasedimentary samples show relatively minor contributions from older mesoto eoarchaean material. the synchronous sedimentation and generation of volcanic material at c. 3.07 ga (mesoarchaean) and at c. 2.83 ga (neoarchaean) suggests there may have been a common tectonic environment operating at both times. when did metamorphism and mineralisation occur? the godthåbsfjord region records a complex history of highgrade metamorphism, which affected different parts of the godthåbsfjord region at different times, from c. 3.8–2.5 ga. the record of particular metamorphic events in different parts of the region has been used to establish the times when different terranes were amalgamated via collisional tectonism (e.g. friend & nutman 2005). the timing of metamorphism and contemporaneous deformation has also been used to constrain the timing of mineralisation events. on storø in central godthåbsfjord (fig. 1) detailed studies of the nature and timing of gold mineralisation have been carried out. here gold-mineralised strataform horizons and cross-cutting quartz veins occur within a neoarchaean supracrustal belt that was deposited at or after 2.8 ga. the goldbearing horizons are deformed by the kilometre-wide, high-grade storø shear zone; thus the generation of the shear fabric post-dates the gold-mineralising event. to place absolute constraints on the timing of gold mineralisation the timing of emplacement of some of the abundant granitic pegmatites that either cross-cut or are deformed by the shear fabric has been investigated (fig. 5). four pegmatites contained 51 supracrustal belts 25 km st or ø 64°n 52°w 51°w 50°w 49°w godthåbsfjord 2.7 2.9 3.1 3.3 3.5 3.7 ga 64°30' n metavolcanic metasedimentary metasedimentary metasedimentary metasedimentary metasedimentary metasedimentary metavolcanic metasedimentary metavolcanic fig. 4. detrital zircon populations linked to the distribution of supracrustal belts. inset shows the location of the main map in greenland. red zircon spectra were collected using the geus la-sf-icp-ms. blue zircon spectra are taken from nutman et al. (2004) and schiøtte et al. (1988). dotted lines show regionally significant age peaks at 3.07, 2.89 and 2.83 ga. 52 igneous zircon crystallised at c. 2.63–2.60 ga. this is the same age as yielded by metamorphic zircons separated from a leucocratic amphibolite, and from metamorphic overgrowths on detrital zircons taken from a melt layer within a migmatised metasedimentary rock, both within the same supracrustal sequence on storø. these results show that a gold-mineralising event occurred within the neoarchaean supracrustal belt during or after sedimentation at c. 2.8 ga and before the formation of the storø shear zone during high-grade metamorphism and partial melting at c. 2.63 ga. these results constrain the timing of mineralisation and may be extended to the investigation of rocks of similar age and setting elsewhere in the region. further work this work forms part of ongoing studies of the supracrustal belts in the godthåbsfjord region carried out with support from the greenland bureau of minerals and petroleum. in addition to the examples given here, the geus la-sf-icpms system can also be used to investigate the cooling history of rocks via u-pb dating of zircon and other minerals, including titanite and monazite. furthermore, by exploring other aspects of zircon geochemistry, such as hf and o isotopes it may also be possible to delve deeper into issues of crustal evolution of the north atlantic craton through the archaean. references garde, a.a. 2005: a mid-archaean island arc complex with gold mineralisation at qussuk, akia terrane, southern west greenland. in: hollis, j.a. (ed.): greenstone belts in the central godthåbsfjord region, southern west greenland: geochemistry, geochronology and petrography arising from 2004 field work, and digital map data. danmarks og grønlands geologiske undersøgelse rapport 2005/42, 215 pp., 1 dvd. hollis, j.a. (ed.) 2005: greenstone belts in the central godthåbsfjord region, southern west greenland: geochemistry, geochronology and petrography arising from 2004 field work, and digital map data. danmarks og grønlands geologiske undersøgelse rapport 2005/42, 215 pp., 1 dvd. frei, d., hollis, j.a., gerdes, a., harlov, d., karlsson, c., vasquez, p., franz, g. & knudsen, c. 2006: advanced in situ geochronological and trace element microanalysis by laser ablation techniques. geological survey of denmark and greenland bulletin 10, 25–28. friend, c.r.l. & nutman, a.p. 2005: new pieces to the archaean terrane jigsaw puzzle in the nuuk region, southern west greenland: steps in transforming a simple insight into a complex regional tectonothermal model. journal of the geological society (london) 162, 147–162. nutman, a.p., friend, c.r.l., barker, s.l.l. & mcgregor, v.r. 2004: inventory and assessment of palaeoarchaean gneiss terraines and detrital zircons in southern west greenland. precambrian research 135, 281–314. schiøtte, l., compston, w. & bridgwater, d. 1988: late archaean ages for the deposition of clastic sediments belonging to the malene supracrustals, southern west greenland: evidence from an ion probe u-pb zircon study. earth and planetary science letters 87, 45–58. authors’ addresses j.a.h., d.f., j.v.g., a.a.g., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: jho@geus.dk m.p., university of copenhagen, øster voldgade 10, dk-1350 copenhagen k. mafic supracrustal amphibolite host rock pegm atite pegm atite pegm atite c. 2.63 ga zircon emplacement age shear zone fabric fig. 5. pegmatite on storø cross-cutting and deformed by the storø shear zone fabric, and neoarchaean supracrustal amphibolite host rock. zircons from the pegmatite indicate it was formed at c. 2.63 ga, placing a minimum age constraint on gold mineralisation in the host rocks. geological survey of denmark and greenland bulletin 15, 2008, 85-88 on 26 december 2004, the eastern part of the indian ocean was hit by a tremendous tsunami created by a submarine earthquake of magnitude 9.1 on the richter scale off the west coast of sumatra. the tsunami also reached the western part of the indian ocean, including the coastal areas of eastern africa. along the coast of kenya (figs 1, 2) it resulted in a sudden increase in water level comparable to a high tide situation. this rather limited consequence was partly due to the great distance to the epicentre of the earthquake, and partly due to the low tide at the time of the impact. hence the reefs that fringe two thirds of the coastline reduced the energy of the tsunami waves and protected the coastal areas. during the spring of 2005, staff members from the geo logical survey of denmark and greenland (geus) carried out field work related to the project kensea – development of a sensitivity atlas for coastal areas of kenya (tychsen 2006; tychsen et al. 2006). local fishermen and authorities often asked what would have been the effect if the tsunami had hit the coastal area during a high tide, and to answer the question geus and the kenya marine and fisheries research institute (kmfri) initiated a tsunami damage projection project. the aim was to provide an important tool for contingency planning by national and local authorities in the implementation of a national early warning strategy. the tsunami damage projection project used the database of coastal resources – kenseabase – that was developed during the kensea project. the topographical maps of kenya at a scale of 1:50 000 have 20 m contour lines, which is insufficient for the tsunami run-up simulation modelling undertaken by the new tsunami project. therefore new sets of aerial photographs were obtained, and new photogrammetric maps with contour lines with an equidistance of 1 m were drawn for a 6–8 km broad coastal zone. the tsunami modelling is based on the assumption that the height of a future tsunami wave would be comparable with the one that reached the coastal area of kenya in december 2004. based on the regional geology of the indian ocean, it appears that the epicentre for a possible future earthquake that could lead to a new tsunami would most likely be situated in the eastern part of the ocean. furthermore, based on a seismological assessment it has been estimated that the largest tsunami that can be expected to reach eastern africa would have a 50% larger amplitude than the 2004 tsunami. it was therefore decided to carry out the simulation modelling with a tsunami wave similar to that of the 2004 event, but with the wave reaching the coast at the highest astronomical tide (scenario 1) and a worst case with a 50% larger 85 kensea – tsunami damage modelling for coastal areas of kenya john tychsen, ole geertz-hansen and frands schjøth © geus, 2008. geological survey of denmark and greenland bulletin 15, 85–88. available at: www.geus.dk/publications/bull mombasa 50 km kenya africa 39°e 40°e 41°e 5°s 4°s 3°s 2°s kenya tanzania lamu gazi indian ocean ungwana bay fig. 4a fig. 3a fig. 3b figs 3c, 4c fig. 4b fig. 1. map of the coastal area of kenya (red frame on index map). the green frames show the locations of the maps shown in figs 3–4. fig. 2. a fishing vessel lying on the beach east of ngomeni. the vessel was wrecked by the tsunami in december 2004 (see fig. 3a for location). amplitude (scenario 2: fig. 3). the 2004 tsunami documented that the coastal belt of mangrove swamps provided some protection to the coastline by reducing the energy of the tsunami. hence we included in this study a scenario 3 (fig. 4), in which the mangrove areas along the coastline were removed. maps for the three scenarios have been produced and show the areas that would be flooded, the degree of flooding, and the distribution of buildings such as schools and hospitals in the flooded areas. in addition, the force and velocity of the wave were calculated (cowi 2006). 86 5 km max. water column 0–1 m 1–2 m 2–3 m 3–4 m 4–5 m 5–6 m 6–7 m 7–8 m 0 metre 39°42′e39°36′e 39°42′e39°36′e 4°0′s 4°6′s 4°0′s 4°6′s waa nyali matuga likoni kisauni mombasa frere town 5 km5 km mida gede watamu mulindi ngomeni malindi 5 km mambrui gongoni marikebuni 40°0′e 40°6′e 3°18′s 3°24′s 3°18′s 3°24′s 40°12′e40°6′e 3°0′s 3°6′s 3°12′s a b c sabaki river fig. 3. scenario 2, showing maximum simulated water levels north of sabaki river, in the watamu-malindi area, and around mombasa for a tsunami reaching the coast at high tide, and with a 50% larger amplitude than the 2004 tsunami. at sea, sea surface elevation is shown relative to mean sea level, whereas on land the water level is relative to the land surface and therefore shows flooding heights (see fig. 1 for location). the run-up simulation model the study used the mike 21 bw model, which is a 2-d hydrodynamic model from the mike modelling suite developed by the danish hydraulic institute. the modelling was undertaken by cowi a/s in denmark (cowi 2006). the model setup included detailed bathymetry and topography of the area and data on the surface properties (e.g. sand, reef, rock, mangrove, forest, town), which provide information on bed resistance. the model covers an area of 49 000 km2 with a grid size of 100 x 100 m. thus the total number of grid cells is 4 900 000. the topographical data mainly derive from the detailed topographic maps that were drawn from the new aerial photographs. where necessary these data were complemented with data extracted from existing topographic maps. the topographical data thus only cover elevations above mean sea level. bathymetrical data were extracted from cmap, a world wide digital navigational chart, by a module that produces bathymetrical data that can be used directly by the mike models. the c-map data include water depths at and below the chart datum. the topography between mean sea level and chart datum was interpolated. no measured boundary data were available, and a normal calibration of the model was therefore not possible. only one single high-resolution time series of water level changes during the 2004 tsunami incident is available, from the port of lamu in the north. the offshore boundary conditions off 87 5 km scenario 3 minus scenario 2 increased water level 0–0.2 m 0.2–0.4 m 0.4–0.6 m 0.6–0.8 m 0.8–1 m >1 m 0 metre mangrove margin 5 km 5 km funzi wasini ramisi shimoni kibuyuni jego vanga nyali likoni kisauni mombasa frere town siyu paté island mbajumali kijingitini chundwa/tundwa 39°18’e 39°24’e39°12’e 4°30’s 4°36’s 4°42’s 39°36’e 39°42’e 39°36’e 39°42’e 4°0’s 4°6’s 4°0’s 4°6’s 41°6’e 2°6’s 2°12’s a b c fig. 4. the effect of mangroves. the figure shows the difference between scenario 3 (all mangrove removed) and scenario 2 (mangrove present) for three important mangrove areas (paté island, the vanga-shimoni-funzi area and the mombasa area). the green 0–0.2 m signature in open water can mostly be regarded as noise from the model. the only marked difference between the models is seen behind areas with dense mangroves (see fig. 1 for location). lamu were back-calculated by a trial and error approach until the model reproduced the recorded time series at lamu. by this approach, bed resistances could not be used for calibration, but were set to well-established values for the various types of sea bed and land surfaces mapped and described by tychsen (2006). the resulting boundary conditions (i.e. the offshore tsunami wave train) were then applied to the full length of the kenyan coast. model results a few examples of the model results are illustrated here, and more details can be found in the project reports (cowi 2006; geus 2007). all data and results are stored in electronic form in the kensea database located at kmfri in kenya. three areas have been selected to give an impression of the effects of a possible new tsunami. the risk of a scenario 2 incident is extremely small, but the model helps to pinpoint the most sensitive areas along the coast (fig. 3). effects of mangrove the mitigating effects of mangrove forests on the destructive powers of tsunamis have been described and discussed by several authors following the december 2004 tsunami (e.g. kathiresan & rejendran 2005). scenario 2 was therefore remodelled with all mangroves removed, i.e. with bed resistance corresponding to normal sea bed (scenario 3). the difference between models with and without mangrove is not obvious, and it was therefore decided to subtract scenario 2 from scenario 3 to isolate the mangrove effects. this method creates some noise and artefacts, but the conclusion is that the effect on water level is small, less than 20 cm in most places; however, behind wide and somewhat exposed mangroves the difference can be up to 60 or 70 cm. in addition, the flooding extends up to 300 m farther inland when the mangrove is removed. the most significant effect is found behind the mangrove areas north of vanga, behind funzi, near gazi, in the south-western part of ungwana bay, and on paté island (fig. 4). the effects in kenya of removing the mangrove are smaller than expected from other studies (gelfenbaum et al. 2007). this is not because the mangroves do not mitigate tsunami waves, but because mangroves along the coast of kenya are mostly found in areas already protected from direct wave exposure, as found in sheltered bays and lagoons, and behind islands or wide reefs. mangroves do not usually grow naturally along the most exposed coastlines characterised by erosion. this is partly due to the exposure, but mainly due to the lack of suitable substrates for the roots. recommendations an important output of the project was the following set of recommendations to the government of kenya: • it is recommended that an indian ocean tsunami warning system is developed, similar to the well-developed system in the pacific ocean. • a post-2004 tsunami study showed a marked lack of knowledge of tsunamis among the coastal communities. hence there is a need to create public awareness of the causes and potential impacts of tsunamis to enable the local population to take appropriate action when an alarm is raised and thus to minimise the effects of future tsunamis. • mangrove forests play an important role in mitigating the impact of tsunami waves. it is recommended that the department of forestry in conjunction with local communities rehabilitate areas where mangrove has been cut down. • in high-risk areas, the provincial administration in collaboration with the disaster and tsunami manage ment committee should educate the population living in those areas. this would help them to cope better with a tsunami disaster, both physically and psychologically. references cowi 2006: tsunami run-up simulation model for the near shore part of the coastal area of kenya, 49 pp. unpublished report, cowi for the geo logical survey of denmark and greenland, copenhagen. gelfenbaum, g., vatvani, d., jaffe, b. & dekker, f. 2007: tsunami inundation and sediment transport in vicinity of coastal mangrove forest. in: kraus, n.c. & rosati, j.d. (eds): coastal sediments ’07. proceedings of the sixth international symposium on coastal engineering and science of coastal sediment processes, 1117–1128. 13–17 may 2007, new or leans, louisiana. geus 2007: the kensea ii project. tsunami damage projection for the coastal area of kenya, 70 pp. unpublished report, geological survey of denmark and greenland, copenhagen. kathiresan, k. & rejendran, n. 2005: coastal mangrove forests mitigated tsunami. estuarine, coastal and shelf processes 65, 601–606. tychsen, j. 2006 (ed.): kensea – environmental sensitivity atlas for coastal area of kenya, 76 pp. copenhagen: geological survey of denmark and greenland. tychsen, j., geertz-hansen, o. & kofoed, j. 2006: kensea – development of an environmental sensitivity atlas for coastal areas of kenya. geo logical survey of denmark and greenland bulletin 10, 65–68. 88 authors’ addresses j.t. & f.s., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: jt@geus.dk o.g.-h., aquasim, slagslunde bygade 37, dk-3660 stenløse, denmark. geological survey of denmark and greenland bulletin 17, 2009, 49-52 in 2008, the geological survey of denmark and greenland began a project in collaboration with the bureau of minerals and petroleum of greenland with the aim to publish a webbased, seamless digital map of the precambrian bedrock between 61°30´ and 64°n in southern west greenland. such a map will be helpful for the mineral exploration industry and for basic research. producing an updated digital map requires additional field work revisiting key localities to collect samples for geochemistry, geochronology and metamorphic petrology. the new data will help us to test and refine existing models and improve general understanding of the geological evolution of the area. here we summarise some results from the 2008 field activities between ame ralik in the north and frederikshåb is blink in the south (fig. 1). the area was mapped in the 1960s and 1970s, and although the 1:100 000-scale maps are of excellent quality, they do not include more recent developments in geochro nology, thermobarometry and geochemistry. a notable exception is the fiske næsset complex (fig. 1), which has re ceived considerable attention after it was first mapped (ellitsgaard-ras mus sen & mouritzen 1954; wind ley et al., 1973; windley & smith, 1974; myers 1985). new tectonic models have been developed since the original 1:100 000 maps were produced, and the tectonic evolution has been com monly ex plained in terms of terrane accretion (friend et al. 1996). friend’s model de fines a number of boundaries that separate terranes of different age and origin and which might have contrasting tectono-metamorphic histories prior to terrane accretion. the current project area includes the northern part and proposed boundary of the tasiusarsuaq terrane, which was amalgamated with the terranes to the north at 2.72 ga, when regional metamorphism affected the region (friend et al. 1996). in addition, windley & garde © geus, 2009. geological survey of denmark and greenland bulletin 17, 49–52. available at: www.geus.dk/publications/bull geological observations in the southern west greenland basement from ameralik to frederikshåb isblink in 2008 nynke keulen, anders scherstén, john c. schumacher, tomas næraa and brian f. windley 49 mainly upper crustal zone lower crustal zone prograde amphibolite facies retrograde amphibolite facies granulite facies quaternary deposits block boundary inferred thrust transition between lower and upper crustal zones late kinematic ttg plutons and granitic rocks orthogneiss metavolcanic belt eoarchaean gneiss anorthosite complex tasiusarsuaq western terrane boundary 64°n 62°n52°w 48° 50 km bjørn esu nd ikk att oq fre der iksh åb isb link tre brødre ter ran e fiskenæsset complex utqqôrqqqq uuuuur tttttuuuuuôrqqutôrqutqutqutqutq tttqôôôqôôôôôqqq tetnitang aggraninnran ttg teg ngran teetei rtalikrtvev alikkkkkkiillrriilllllliiiiii talikkee ttavertaaerta kkkki t li qôrqut granite buks efjo rde n amera lik grædef jord ikk att up nunaa sermilik bjørnesund kvanefjord nuuk paamiut qarli it n unaat th rust qilan ngaa rsu it qilan ngaa ssu a tasiusarsuaq terrane majo qqa p q aav a inland ice ilivertalik granite davis strait greenland fig. 1. simplified geological map of the central part of the north atlantic craton in southern west greenland showing the main rock com ponents, patterns of metamorphic facies and three crustal blocks (modified after windley & garde 2009). rosa_2008:rosa-2008 01/07/09 15:48 side 49 (2009) proposed a model in which a series of blocks represent crustal sections that (except for the sermilik block; fig. 1) display a systematic metamorphic progression from amphibolite facies to granulite facies rocks from south to north. each block represents a number of island and continental arcs that were amalgamated by collision into the growing north atlantic craton. in their model, the project area consists of the sermilik, bjørnesund and kvanefjord blocks, where each block represents a combination of lateral and vertical crustal growth. however, contrary to the terrane model, two blocks can have a common origin. quartzo-feldspathic rocks grey tonalite-trondhjemite-granodiorite-type gneiss es (ttg) have intrusive ages of 2.92–2.84 ga that represent the main crust-generating period (schiøtte et al. 1989; næraa & scher stén 2008). there is growing evidence from zircon u-pb-hf geochronology for an older previously unrecognised meso archaean crustal component within the tasiusarsuaq terrane (næraa & scherstén, unpublished data). the extent of the older component is currently unclear, but it is apparently most widespread in the northern part of the terrane. in addition, a number of strongly deformed and veined tonalitic gneisses of presumed early neoarchaean age are younger, about 2.72 ga, and formed from older pre-existing crust. granite (sensu lato) intrusions are variably deformed but post-date the majority of the grey gneisses. the main granite, the ilivertalik augen granite, has feldspar megacrysts, and is in some places orthopyroxene-bearing and has been dated to 2.8 ga (pidgeon & kalsbeek 1976; næraa & scherstén, unpublished data). the main volume of this granite intrudes the sermilik block, but satellite intrusions, including the type locality at ilivertalik mountain, are found in the bjørnesund block, implying that these blocks were a single crustal unit by 2.8 ga. at ilivertalik mountain, kalsbeek & myers (1973) suggested that the intrusion was charnockitic and was emplaced under granulite-facies conditions. mica schists, commonly with garnet or sillimanite, occur throughout the area. these were originally mapped as having sedimentary protoliths. to our knowledge, all recognised structures in these rocks are tectono-metamorphic, and their supracrustal origin seems to have been mostly inferred from their common relationship with amphibolites or their aluminium-rich or quartzitic compositions. amphibolites and metagabbro-anorthosites supracrustal rocks with a range of compositions occur throughout the area, but are dominated by amphibolites. preservation varies, with readily identified primary textures in 50 temperature (°c) 400 500 600 700 800 900 1000 1100 1200 1 0 2 3 4 5 6 7 8 9 10 p (k ba r) d epth (km ) calculated geotherms facies boundaries reaction boundaries metamorphic conditions consistent with field observations in 2008 possible p-t trajectory of a plagioclase lherzolite boudin (fig. 3) reaction of olivine+plagioclase to orthoand clinopyroxene in plagioclase lherzolite boudin am gn sil ky and melting in quartzfeldspar rocks melting of ultramafic rocks qz m us sil kfs opx cpx ol plg bio qz opx kfs l t spinel lhz plagioclase lhz ~2.7 ga ~3.2 ga ~2.56 ga ~3.0 ga 3 2 4 1 eb ea su dil os eti na rg de ar ut as -l a gs 35 25 15 5 fig. 2. ranges of p–t estimates from previous work by wells (1979), griffin et al. (1980) and riciputi (1990); geotherms based on measured and estimated contents as a function of time of radiogenic elements in basaltic and felsic green land rocks using methods of kamber et al. (2005). the calculations are based on a crustal thickness of 40 km and a two-layer model with 10 km of mafic rocks at the base of the crust overlain by 30 km of felsic rocks. depths in kilometres are based on an average crustal density of 2.75 g/cm3. parameters: man tle heat flow: 20 mw/m2; mafic and felsic rockheat production are: 0.262 and 3.478 µw/m3 at 3.2 ga (possible maximum age for this part of the terrane), 0.244 and 3.254 µw/m3 at 3.0 ga, 0.220 and 2.945 µw/m3 at 2.7 ga, 0.210 and 2.811 µw/m3 at 2.56 ga (age of the qôrqut granite). the numbered blue fields represent the ranges of p–t estimates: 1 = sermilik (griffin et al. 1980); 2 = sermilik and qilan ngaassua (riciputi et al. 1990); 3 = sermilik (wells 1979); 4 = qilanngaarsuit (wells 1979). am, amphibolite facies; and, andalusite; bio, biotite; cpx, clinopyroxene; ea, epidoteamphibolite facies; eb, epidote-blueschist facies; gn, granulite facies; gs, greenschist facies; kfs, k-feldspar; ky, kyanite; l, liquid; lhz, lherzolite; mus, muscovite; ol, olivine; opx, orthopyroxene; plg, plagioclase; qz, quartz; sil, sillimanite. rosa_2008:rosa-2008 01/07/09 15:48 side 50 some areas and amphibolite lenses in others. the ikkattup nunaa belt on the islands in the ikkattoq fjord just north of frederikshåb isblink (fig. 1) is one of the best-preserved volcanic belts in the region (andersen & friend’s 1973 ravns storø belt). pillow lavas and volcanic bombs are well preserved here, and these rocks likely formed in shallow water with partly explosive volcanism. rocks with calc-alkaline basaltic to andesitic compositions predominate (k. szilas et al., unpublished data 2009), and analytical results suggest a convergent margin setting with an age of 2.91 ± 0.01 ga (nutman et al. 2004). in 2008 we discovered that some major amphibolite belts at majoqqap qaava (fig. 1) contain two components, namely lithic tuff-dominated, metavolcanic rocks and massive homogeneous amphibolite sheets, which are similar to the two main components of the ikkattup nunaa belt, thus a common arc origin is likely. the fiskenæsset complex is a layered igneous complex containing ultramafic rocks, gabbros, leucogabbros, anorthosites and chromitites, which formed largely by cumulate processes (myers 1985). in 2008, we established that the lower main ultramafic unit, best exposed at majoqqap qaava, consists of a succession of layered dunites intruded by a major sill complex that comprises more than 25 clinopyroxene-hornblende sills, some of which are up to 10 m thick. the sills contain xenoliths of dunite, send apophyses into adjacent dunites and have discordant contacts against layered dunites. this observation means that the early history of the fiskenæsset complex involved intrusion of a second, hydrous magma batch into earlier crystallised, olivine-rich cumulates. meta-gabbros and meta-anorthosites of the fiskenæsset complex are closely associated with amphibolites of supra crustal origin, and they could be part of the same magmatic system. escher & myers (1975) believed that the fiskenæsset complex was intrusive into the amphibolites. however, the observed contacts are tectonic, and direct evidence for primary intrusive relationships is lacking. clearly, discordant anor thosite dykes cross-cut strongly deformed metagabbros within the complex (e.g. at 63.1156°n, 50.7155°w). these dykes consist of plagioclase and orthopyroxene and likely post-date all deformation. tectono-metamorphic development the qarliit nunaat thrust forms the boundary between the tre brødre and tasiusarsuaq terranes (fig. 1; friend et al. 1996). south of buksefjorden there is a very high-strain, high-grade shear zone several kilometres wide between the færingehavn and the tasiusarsuaq terranes (stainforth 1977; crowley 2002). however, on either side of this shear zone, the tectono-metamorphic styles are different; to the north of buksefjorden we found no evidence for a metamorphic or structural discontinuity. here, rocks previously considered to be part of the tre brødre terrane might therefore be part of the tasiusarsuaq terrane. an abrupt change in metamorphic grade, from amphibolite facies just south of grædefjord to granulite facies a little farther south, within the tasiusarsuaq terrane, is associated with intensive shearing as predicted in the crustal block model of windley & garde (2009). confirmation of the model comes from the fact that in 2008 we discovered that the boundary is occupied by a north-dipping, over 250-mwide shear zone with a down-dip lineation, which contains augen gneisses, cataclasites and mylonites. the boundary separates amphibolite facies gneiss with granulite facies relicts to the south from prograde amphibolite facies gneisses to the north, as also predicted by kalsbeek (1976). pressure –temperature (p–t) estimates are sparse in the present field area and focused on pyroxene-garnet assemblages (amphibolites) that are restricted to upper amphibolite and granulite facies. figure 2 shows the locations of estimated geotherms for a 40 km thick crust. the geotherms that bracket the (3.0–2.7 ga) metamorphism suggest pressures of 5–6 kbar for the conditions of the amphibolite facies to granulite facies transition. these p–t-t (t = time) conditions are consistent with field observations, e.g. the only aluminosilicate phase found in 2008 was sillimanite, and the observed peak assemblages are consistent with conditions near the amphibolite facies and granulite facies transition at intermediate pressure. granulite-facies rocks occur more rarely than suggested on the current 1:100 000 scale maps, although, in the area around the fiskenæsset complex granulite-facies rocks are 51 1 cm fig. 3. spinel-bearing plagioclase lherzolite showing the breakdown of olivine and plagioclase to form coronas of clinopyroxene and orthopyroxene. the stability limits of the plagioclase lherzolite are below c. 9 kbar and 1200ºc, and the reaction would have occurred between about 700°c, 4 kbar and 870ºc, 7 kbar. location at 63.932ºn, 51.219ºw. rosa_2008:rosa-2008 01/07/09 15:48 side 51 52 abundant. some of the area designated as granulite facies contains amphibolite-facies assemblages, but is interpreted to have once attained granulite-facies conditions. however, part of the rocks previously considered retrograde from granulitefacies rocks might better be interpreted as prograde amphibolite-facies rocks that never reached granulite-facies con ditions. full analysis of the p–t–t trajectories will require detailed textural and chemical analysis of these rocks. garnet-bearing mica schists and amphibolites from ame ralik fjord and frederikshåb isblink suggest garnet growth at the expense of plagioclase, which is consistent with a metamorphic event dominated by pressure increase. these observations are best explained by thrust tectonics. evidence for isobaric or near-isobaric cooling was observed in reaction textures of the plagioclase lherzolite from north of buksefjorden described in fig. 3. the lherzolite occurs as a boudin within a refolded zone of amphibolite and ultramafic rocks in grey gneiss. ultramafic boudins are associated with pegmatite dykes, and this resulted in alteration of the ultramafic rocks. similar corona textures were noted at majoqqap qaava and were interpreted as late igneous or early metamorphic reactions by myers & platt (1977). cataclastic structures are abundant in the area between ameralik and frederikshåb isblink (e.g. stainforth 1977). brittle deformation occurs as an overprint of earlier ductile structures. pseudotachylytes were observed at a few localities, although cohesive fault rocks (cataclasites) are more common. references andersen, l.s. & friend, c. 1973: structure of the ravns storø amphibolite belt in the fiskenæsset region. rapport grønlands geologiske undersøgelse 51, 37–40. crowley, j.l. 2002: testing the model of late archaean terrane accretion in southern west greenland: a comparison of timing of geological events across the qarliit nunaat fault, buksefjorden region. pre cam brian research 116, 57–79. ellitsgaard-rasmussen, k. & mouritzen, m. 1954: an anorthosite occurrence from west greenland. meddelelser fra dansk geologisk for ening 12, 436–442. escher, j.c. & myers, j.s. 1975: new evidence concerning the original relationships of early precambrian volcanics and anorthosites in the fiskenæsset region, southern west greenland. rapport grønlands geologiske undersøgelse 75, 72–76. friend, c.r.l., nutman, a.p., baadsgaard, h., kinny, p.d. & mcgregor, v.r. 1996: timing of late archaean terrane assembly, crustal thickening and granite emplacement in the nuuk region, southern west green land. earth and planetary science letters 142, 353–365. griffin, w.l., mcgregor, v.r., nutman, a.p., taylor, p. n. & bridgwater, d. 1980: early archaean granulite-facies metamorphism south of ame ralik, west greenland. earth and planetary science letters 50, 59– 74. kalsbeek, f. 1976: metamorphism in the fiskenæsset region. rapport grønlands geologiske undersøgelse 73, 34–41. kalsbeek, f & myers, j.s. 1973: the geology of the fiskenæsset region. rapport grønlands geologiske undersøgelse 51, 5–18. kamber, b.s., whitehouse, m.j., bolhar, r. & moorbath, s. 2005: volcanic resurfacing and the early terrestrial crust: zircon u–pb and ree constraints from the isua greenstone belt, southern west greenland. earth and planetary science letters 240, 276–290. myers, j.s. 1985: stratigraphy and structure of the fiskenæsset complex, southern west greenland. bulletin grønlands geologiske under søgelse 150, 72 pp. myers, j.s. & platt, r.g., 1977: mineral chemistry of layered archaean anorthosite at majorqap qâva, near fiskenæsset, southwest greenland. lithos 11, 59–72. næraa, t. & scherstén, a. 2008: new zircon ages from the tasiusarsuaq terrane, southern west greenland. geological survey of denmark and greenland bulletin 15, 73-76. nutman, a.p., friend, c.r.l., barker, s.l.l. & mcgregor, v.r. 2004: inventory and assessment of palaeoarchaean gneiss terrains and detrital zircons in southern west greenland. precambrian research 135, 281–314. pidgeon, r.t. & kalsbeek, f. 1978: dating of igneous and metamorphic events in the fiskenaesset region of southern west greenland. cana dian journal of earth sciences 15, 2021–2025. riciputi, l.r., valley, j.w. & mcgregor, v.r. 1990: conditions of archean granulite metamorphism in the godthab-fiskenaesset region, southern west greenland. journal of metamorphic geology 8, 171–190. schiøtte, l., compston, w. & bridgwater, d. 1989: u-pb single-zircon age for the tinissaq gneiss of southern west greenland: a controversy resolved. chemical geology 79, 21–30. stainforth, j.g. 1977: the structural geology of the area between ameralik and buksefjorden, southern west greenland, 480 pp. unpublished ph.d. thesis, exeter university, uk. wells, p.r.a. 1979: chemical and thermal evolution of archaean sialic crust, southern west greenland. journal of petrology 20, 187–226. windley, b.f. & garde, a.a. 2009: arc-generated blocks with crustal sections in the north atlantic craton of west greenland: crustal growth in the archean with modern analogues. earth-science reviews 93, 1–30. windley, b.f. & smith, j.v. 1974: the fiskenæsset complex, west green land, part ii. general mineral chemistry from qeqertarssuatsiaq. bulle tin grønlands geologiske undersøgelse 108, 54 pp. windley, b.f., herd, r.k. & bowden, a.a. 1973: the fiskenæsset complex, west greenland, part i. a preliminary study of the stratigraphy, petrology, and whole rock chemistry from qeqertarssuatsiaq. bulletin grønlands geologiske undersøgelse 106, 80 pp. authors’ addresses n.k., a.s.1 & t.n., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ntk@geus.dk j.c.s., department of earth sciences, university of bristol, bristol bs8 1rj, uk. b.f.w., department of geology, university of leicester, leicester le1 7rh, uk. 1present address: department of geology, lund university , sölvegatan 12, s-223 62 lund, sweden. rosa_2008:rosa-2008 01/07/09 15:48 side 52 geological survey of denmark and greenland bulletin 11, 53-60 53© geus, 2006. geological survey of denmark and greenland bulletin 11, 53–60. available at: www.geus.dk/publications/bull a lead isotope study of an archaean gold prospect in the attu region, nagssugtoqidian orogen, west greenland henrik stendal, robert frei and bo møller stensgaard this paper presents a lead isotope investigation of a gold prospect south of the village attu in the northern part of the nagssugtoqidian orogen in central west greenland. the attu gold prospect is a replacement gold occurrence, related to a shear/mylonite zone along a contact between orthogneiss and amphibolite within the nagssugtoqidian orogenic belt. the mineral occurrence is small, less than 0.5 m wide, and can be followed along strike for several hundred metres. the mineral assemblage is pyrite, chalcopyrite, magnetite and gold. the host rocks to the gold prospect are granulite facies ‘brown gneisses’ and amphibolites. pb-isotopic data on magnetite from the host rocks yield an isochron in a 207pb/204pb vs. 206pb/204pb diagram, giving a date of 3162 ± 43 ma (mswd = 0.5). this date is interpreted to represent the age of the rocks in question, and is older than dates obtained from rocks elsewhere within the nagssugtoqidian orogen. pb-isotopic data on cataclastic magnetite from the shear zone lie close to this isochron, indicating a similar origin. the pb-isotopic compositions of the ore minerals are similar to those previously obtained from the close-by ~2650 ma rifkol granite, and suggest a genetic link between the emplacement of this granite and the formation of the ore minerals in the shear/mylonite zone. consequently, the age of the gold mineralisation is interpreted to be late archaean. keywords: archaean, gold, magnetite, pb isotopes, geochronology, west greenland ________________________________________________________________________________________________________________________________________________________________________________________ h.s. & b.m.s., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: hst@geus.dk r.f., geological institute, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. discovery of the gold prospect described in this study was due to the find of a mineralised sample, which karl markussen from attu submitted to the bureau of minerals and petroleum in greenland. the geological survey of denmark and greenland (geus) visited the locality in 2001 and in 2002 (stendal et al. 2002, 2004), and the present paper reports pb-isotopic data for minerals from the prospect and its surroundings. the attu gold prospect lies within the nagssugtoqidian orogen of west greenland (fig. 1), where geological mapping and exploration has been carried out for decades by the geological survey, the danish lithosphere centre, university research groups and exploration companies (e.g. kalsbeek et al. 1987; connelly et al. 2000; van gool et al. 2002). in addition to the general investigations, steenfelt (2001) has summarised geochemical signatures from stream sediments, rasmussen & van gool (2000) have described geophysical aspects, and steenfelt et al. (2002), stendal & schønwandt (2003) and stendal et al. (2004) have described mineral occurrences and their economic potential. an overview of the mineral occurrences in the entire region has been presented by stendal et al. (2004). detailed, mainly zircon u-pb geochronological data from the nagssugtoqidian orogen have been presented by kalsbeek & nutman (1996), connelly & mengel (2000) and connelly et al. (2000), and pb-pb, rb-sr and sm54 nd whole-rock isotope data from the region have been reported by kalsbeek et al. (1984, 1987), taylor & kalsbeek (1990) and whitehouse et al. (1998). in addition, some pb-isotopic work has been carried out on sulphide separates, mainly pyrite, from mineral occurrences in the disko bugt region (stendal 1998). geological setting the palaeoproterozoic nagssugtoqidian orogen of west greenland (van gool et al. 2002) is located between the archaean north atlantic craton to the south and a lesserknown continental mass to the north that includes the palaeoproterozoic rinkian fold belt. most of the orogen consists of variably reworked archaean orthogneisses. several thin belts of supracrustal and intrusive igneous rocks occur within this gneiss terrain. granitoid rocks and numerous pegmatites intrude the gneisses. formations of palaeoproterozoic age are limited to the arfersiorfik and sisimiut igneous suites and minor supracrustal sequences (connelly et al. 2000). the attu area itself is located in the southern part of the northern nagssugtoqidian orogen (nno; fig. 1). the metamorphic grade is granulite facies; metamorphism and deformation of the archaean granitoid rocks in the nno gradually decrease northwards, from granulite to amphibolite facies, and from high strain to lower strain with more open structures. steeply and shallowly dipping shear and fault zones are common in contact zones between 68° inland ice 54°54°54° 67°30' attuattu gold prospect ar nssz nno cno rifkol undifferentiated orthogneiss attu gold prospect amphibolite and metasedimentary rocks archaean mainly metasedimentary rocks palaeoproterozoic a 18 60 74 18 64 74 55 58 60 60 56 70 62 70 48 70 63 70 62 70 62 82 50 60 15 40 75 1 km attu gold prospect 53 °3 0' 67°51' 67°48' 53 °2 1' biotite gneiss, mainly granodioritic-tonalitic orthopyroxene gneiss, mainly tonalitic granite/charnockite amphibolite metasedimentary and metavolcanic rocks strike and dip of dominant lithological layering and foliation direction and plunge of fold axis, measured, constructed rock samples with elevated gold values fault, shear zone, mylonite zone quaternary deposits kangiussaq tateraat 50 km50 km50 km 50°52° b 500 km inland ice greenland inland ice greenland iceland ketilidian orogen north atlantic craton nagssugtoqidian orogen rinkian fold belt canada fig. 1. a: geological map of the attu region with index map of greenland. cno, central nagssugtoqidian orogen; nno, northern nagssugtoqidian orogen; nssz, nordre strømfjord shear zone. b: geological map of the attu gold prospect area (modified from olesen 1984). 55 different lithologies. major fault zones generally strike nne to ne. the major nordre strømfjord shear zone (van gool 2002) is located c. 20 km south of the study area. the shear zone is traceable from the coast to the inland ice and forms the southern boundary of the nno. the gneisses of the nno are late archaean, with ages between 2870 and 2700 ma (kalsbeek & nutman 1996; connelly & mengel 2000; hollis et al. 2006, this volume; thrane & connelly 2006, this volume). discordant sheets of granitoid rocks of archaean age occur in the centre of the nno and large charnockite/granite bodies including the rifkol granite are situated 20 km to the northwest and just south of the study area (fig. 1; hansen 1979; kalsbeek et al. 1984). only a few younger palaeoproterozoic ages have been obtained from the nno: thrane & connelly (2006, this volume) have obtained an approximate depositional age of the naternaq supracrustal belt some 80 km north-east of attu of c. 1950–1900 ma, and an undeformed pegmatite between attu and aasiaat has yielded an age of c. 1790 ma (connelly & mengel 2000). the attu gold prospect the attu gold prospect is located south of the village attu within a 100–330 m wide, complex tract hosting several parallel shear/mylonite zones and faults that strike nne to ne and dip 60–70°w (figs 1, 2). the fault zone can be followed along strike in a north-easterly direction for several kilometres. the host rocks are layered, brown 2.24 ppm au 0.34% cu 124 ppb au 0.11% cu 12 ppb au 17 ppb au 2–6 ppm au < 0.5% cu fig. 2. the site of the attu gold prospect. gold values are given for rock samples. red circles are sample sites with gold values obtained from fine-grained stream sediments (see fig. 1 for geographical location). the red line shows the approximate position of the gold-bearing zone in the shear/mylonite zone. the black line shows the approximate position of a stream. fig. 3. layered brown gneiss with black bands of amphibolite. the hammer shaft is 50 cm long. 56 weathering gneiss and amphibolite (fig. 3). at the western border of the tract a gold-bearing shear/mylonite zone follows the contact between brown gneisses and amphibolites. the gold-bearing shear/mylonite zone (fig. 4) is invaded by pegmatite sheets as well as centimetre-wide veins consisting of red alkali-feldspar and quartz with occasional pyrite and magnetite. the estimated relative volume of pegmatite in the tract varies from 1 to 10% (stendal et al. 2002, 2004). the most promising gold showings are found in a coastal profile along the shear/mylonite zone, which can be followed along strike for several hundreds of metres (figs 1, 2). the studied site is a cliff exposure consisting of mylonite (fig. 4) and a rusty weathered band (10–20 cm in width) mineralised with pyrite, magnetite and some chalcopyrite (fig. 5). pyrite and chalcopyrite replace magnetite. the magnetite is predominantly cataclastic in nature, but recrystallised ore also occurs. the gold is found within pyrite and chalcopyrite. the gangue mineralogy comprises quartz, k-feldspar, muscovite, biotite and carbonates (calcite, dolomite and/or ankerite). the mylonite zone is silicified at the contact with the mineralised zone, and sulphide-rich parts are weathered. secondary goethite and malachite are common (fig. 5). the ore is structurally controlled by and confined to favourable sites (sulphide-bearing zones) within the mylonite/ shear/fault zone. the attu gold prospect has returned reproducable gold fig. 4. the gold bearing mylonite zone. kfeldspar occurs on the right side of the yellow magnet pen (10 cm long). the zone also contains pyrite, chalcopyrite and magnetite. amphibolite and orthogneiss (host rocks) 446601 magnetite 14.631 0.007 14.642 0.009 44.688 0.033 0.961 0.939 446602 magnetite 15.051 0.014 14.752 0.015 36.702 0.040 0.969 0.942 446610 magnetite 17.540 0.051 15.361 0.046 37.613 0.112 0.977 0.988 446614 magnetite 17.002 0.025 15.225 0.024 38.086 0.061 0.976 0.967 shear zone and mineralised rock 446616 magnetite 15.423 0.023 14.844 0.023 41.598 0.068 0.979 0.957 2000368 magnetite 15.286 0.009 14.832 0.010 41.201 0.034 0.962 0.936 481093 magnetite 14.247 0.042 14.625 0.044 41.821 0.130 0.982 0.960 446615 pyrite 14.241 0.009 14.587 0.010 42.001 0.035 0.963 0.934 481078 pyrite 14.447 0.011 14.633 0.012 40.805 0.039 0.967 0.925 446616 k-feldspar 15.123 0.008 14.893 0.010 36.451 0.029 0.958 0.932 table 1. pb-isotopic ratios of magnetite, pyrite and k-feldspar from the attu gold prospect and its host rocks sample number mineral 206pb/204pb ± 2σ* 207pb/204pb ± 2σ 208pb/204pb ± 2σ r1** r2† * errors are two standard deviations absolute (ludwig 1990). ** r1 = 206pb/204pb versus 207pb/204pb error correlation (ludwig 1990). † r2 = 206pb/204pb versus 208pb/204pb error correlation (ludwig 1990). 57 concentrations in the range 2.3–5.8 ppm. other localities in the same fault structure yielded 2.24 ppm and 124 ppb au (fig. 2). the gold concentrations are positively correlated with concentrations of copper, and gold-bearing samples often contain magnetite. two stream sediment samples yielded anomalous gold concentrations of 12 ppb and 17 ppb au, respectively (fig. 2). the host gneisses are brownish in colour and comprise orthopyroxene, amphibole, biotite and feldspar, but little quartz. magnetite is in equilibrium with the rock forming minerals and has the same granular texture. in the amphibolites magnetite forms up to millimetre-thick layers, and also occurs in disseminated form. within the mylonite zone, magnetite occurs as a primary phase in the host gneiss as cataclastic grains with cracks filled with pyrite and chalcopyrite, and as a residual phase resulting from sulphide replacement. ten samples were analysed for pb-isotopic compositions. fig. 5. the gold bearing zone (10 cm wide) within the mylonite zone, with malachite and rusty weathered sulphides. 14 15 16 17 1918 14.4 14.6 14.8 15.0 15.2 15.4 15.6 15.8 206pb/204pb 20 7 p b/ 20 4 p b 2400 2000 1600 1200 800 400 0 kfsp mt mt mt mt brown gneiss py py mt amphibolite mt amphibolite mt amphibolite rifkol granite trend line c. 2650 ma 60 ma lower intercept 3143 ma age = 3162 ± 43 ma mswd = 0.50 fig. 6. 206pb/204pb–207pb/204pb diagram for minerals from the attu area listed in table 1. open squares, mineral separates from the gold-bearing mylonite zone. red squares, host rock data. black diamonds, whole-rock samples from the rifkol granite for comparison (data from kalsbeek et al. 1984). mt, magnetite; py, pyrite; kfsp, k-feldspar. blue curve, the pb-isotopic growth curve from stacey & kramers (1975). 58 analytical methods the pb-isotopic study was carried out on magnetite from host gneisses and amphibolites, and on k-feldspar, magnetite and pyrite from the shear zone-hosted mineralised zone (table 1). the isotope analyses were carried out at the danish centre for isotope geology, geological institute, university of copenhagen. near-pure mineral fractions were separated from dry split aliquots of crushed and sieved (100–200 µm) rock powders using a hand magnet, a frantz isodynamic separator and heavy liquid techniques. ore minerals were dissolved in concentrated aqua regia. total procedural blanks for pb amounted to < 120 pg which is considered insignificant for the measured pb-isotopic results, relative to the amount of sample pb estimated from the mass spectrometer signal intensities. isotope analyses were carried out on a vg sector 54-it instrument in static collection mode. fractionation for pb was controlled by repetitive analysis of the nbs 981 standard (values of todt et al. 1993) and amounted to 0.103 ± 0.007% / amu (2 σ; n = 11). all results are quoted with 2 σ precisions. results the pb-isotopic compositions of mineral separates from the gold-bearing mylonite zone and its host rocks are listed in table 1. in the 207pb/204pb vs. 206pb/204pb diagram shown in fig. 6, the pb-isotopic compositions of magnetite from the four host rock samples of brown gneiss and amphibolite (red squares) define a straight line with a slope corresponding to 3162 ± 43 ma (mswd = 0.50). this line has intercepts with the stacey & kramers (1975) pb-isotopic growth curve at 3143 and 60 ma. based on the good fit of the data points on the isochron, and the agreement of the isochron age with the intercepts of the growth curve, we interpret the 3162 ma date as the age of the rocks in question. however, farther south, in the central part of the nagssugtoqidian orogen, palaeoproterozoic granulite facies metamorphism has led to u loss in archaean rocks, resulting in pb-isotopic compositions plotting above and to the left of an 2800 ma reference isochron (whitehouse et. al. 1998). if this process had also taken place in the area of the present study, the 3162 ma date might give a false impression of the age of the rocks. however, the good fit of the data points on the isochron and the agreement of the intercepts with the stacey & kramers (1975) growth curve with the isochron age would then be accidental, a coincidence which we regard as very unlikely. six mineral separates from the gold-bearing mylonite zone (fig. 6, open squares) lie close to or slightly above the isochron obtained for magnetite from the host rocks. the most primitive 206pb/204pb and 207pb/204pb ratios have been measured in pyrite and magnetite from the ore-bearing zone, whereas the two primary magnetites with cataclastic texture from within the shear zone plot very close to the host rock magnetite isochron. this suggests that their crystallisation took place at about the same time as the magnetites from outside the shear zone. whole-rock pb-isotopic ratios from the rifkol granite (kalsbeek et al. 1984) are also plotted on fig. 6 for comparison, and the isotopic values are listed in table 1. the errorchron defined by these samples has a slope corresponding to an age of 2653 ± 110 ma, which has been interpreted as emplacement age of the granite (kalsbeek et al. 1984). this errorchron is oblique and discordant to the isochron obtained for magnetite from the host rocks, but the three least radiogenic data points from ore minerals associated with native gold from within the shear zone are conformable with this younger trend. this suggests that the fluids in the shear zone from which the gold mineralisation was deposited were somehow genetically linked to the intrusion of the rifkol granite. alkali feldspar from the shear zone has its own pb-isotopic signature, which is neither compatible with a ‘rifkol’ source nor with a source typical of the immediate host rocks. the uranogenic vs. thorogenic isotopic pattern (not shown in a figure) is more disperse than the uranogenic pattern and does not add to the understanding of the uranogenic pb-isotopic data; as expected, it mostly reflects the differences in u and th concentrations in the different analysed minerals. summary, discussion and conclusions the attu gold prospect is small. the gold mineralised zone does not exceed 0.5 m in width, and its length is now known to be only a few hundred metres. gold has also been detected along strike several kilometres away, but the mineralisation does not show a continuous outcrop pattern. however, the fact that gold is present indicates that the ne-striking shear/mylonite zone is mineralised and that hydrothermal activity seems to have occurred in most of the prominent lineaments in the region. the goldbearing sulphide deposit is of replacement type, where pyrite and chalcopyrite grew at the expense of e.g. magnetite. it is envisaged that gold was introduced contemporaneously with the replacement processes. reworked archaean orthogneisses dominate all segments of the nagssugtoqidian orogen. published age deter59 minations range from 2870–2700 ma (e.g. kalsbeek & nutman 1996; connelly & mengel 2000), but no chronological information has yet been available from the attu region. the 3162 ± 43 ma magnetite age obtained from the attu host rocks suggests that the rocks in this part of the nagssugtoqidian orogen may be significantly older than similar rocks elsewhere in the orogen. however, smnd isotope data from archaean gneisses in the central part of the orogen suggest the involvement of pre-2800 ma crustal material (possibly 3100 ma or older) in their source (whitehouse et al. 1998). large parts of the nagssugtoqidian orogen underwent palaeoproterozoic granulite facies metamorphism around 1850 ma (e.g. willigers et al. 2001), which resulted in severe disturbance of the pbisotopic evolution of the rocks (whitehouse et al. 1998). in view of the well-preserved 3162 ma isochron relationships for the attu gneisses it appears possible that these rocks escaped high-grade nagssugtoqidian metamorphism and that granulite facies metamorphism here is of archaean age, in agreement with the conclusions of mazur et al. (2006, this volume) and thrane & connelly (2006, this volume). the pb-isotopic data of the gold bearing samples (fig. 6) suggest a genetic link between the rifkol granite intrusion and the fluids percolating through the shear zone, implying an archaean age of the mineralisation. without further analytical work we are unable to elaborate and comment on a possible source of the pb that has been incorporated into the k-feldspar in the shear zone. acknowledgements the authors acknowledge f. kalsbeek, p.m. holm and an anonymous reviewer for thorough criticism and constructive suggestions, which greatly improved the manuscript. the authors would also like to thank the participants in the resource assessment programme mineral resources of the precambrian shield of central west greenland (66°– 70°15′n) for valuable discussions concerning mineralising events in the region. input from other scientists in the region is also gratefully acknowledged. special thanks go to karl markussen, attu, who submitted the first gold bearing sample, for showing us the exact sample locality. references connelly, j.n. & mengel, f.c. 2000: evolution of archean components in the paleoproterozoic nagssugtoqidian orogen, west greenland. geological society of america bulletin 112, 747–763. connelly, j.n., van gool, j.a.m. & mengel, f.c. 2000: temporal evolution of a deeply eroded orogen: the nagssugtoqidian orogen, west greenland. canadian journal of earth sciences 37, 1121–1142. hansen, b.f. 1979: some charnockitic rocks in the nagssugtoqidian of west greenland. rapport grønlands geologiske undersøgelse 89, 85–96. hollis, j.a., keiding, m., stensgaard, b.m., van gool, j.a.m. & garde, a.a. 2006: evolution of neoarchaean supracrustal belts at the northern margin of the north atlantic craton, west greenland. in: garde, a.a. & kalsbeek, f. 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journal of petrology 42, 1729– 1749. __________________________________________________________________________________________________________________________________________________________________________________ manuscript received 4 october 2004; revision accepted 19 december 2005 research article stokholm et al. 2021: geus bulletin 47. 5369. https://doi.org/10.34194/geusb.v47.5369 1 of 8 validation of airborne and satellite altimetry data by arctic trucks citizen science andreas stokholm1 , sine m. hvidegaard1 , rene forsberg1 , sebastian b. simonsen1* 1dtu space, national space institute, technical university of denmark, lyngby, denmark 1. introduction the launch of the european space agency’s (esa’s) first european remote sensing satellite (ers-1) in 1991 made the greenland ice-sheet-wide monitoring of the surface-elevation change possible. since then, satellite altimeters have provided an unbroken record of changes in ice-sheet surface elevation (forsberg et al. 2017; shepherd et al. 2018 sørensen et al. 2018a). this 30-year record from ers-1, ers-2, envisat, icesat, cryosat-2, and the latest icesat-2 satellites will continue into the future with the commissioning of the sentinel-3 satellite series by the european commission (seitz et al. 2010). this record has proven *correspondence: ssim@space.dtu.dk received: 02 jun 2020 accepted: 16 mar 2021 published: 28 may 2021 keywords: satellite validation, altimetry, cryosat-2, operation icebridge, citizen science abbreviations: als: airborne lidar scanning atm: airborne topographic mapper cryovex: cryosat-2 validation experiment dem: digital elevation model dtu space: national space institute, technical university of denmark esa: european space agency gimp: greenland ice mapping project gnss: global navigation satellite system gris: greenland ice sheet lrm: low resolution mode parca: program for arctic regional climate assessment ppp: precise point positioning promice: programme for monitoring of the greenland ice sheet rmse: root mean square error rtk: real-time kinematic sarin and sin: sar interferometric mode siral: synthetic aperture radar (sar) interferometric radar altimeter. geus bulletin is an open access, peerreviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: signe hillerup larsen (geus, denmark) reviewed by: laurence gray (university of ottawa, canada) and one anonymous reviewer. funding: none declared competing interests: none declared additional files: none provided abstract the elevation of ice sheets changes due to climate change, and satellite altimetry is the preferred tool for measuring ice sheet-wide height changes. in situ validation is needed to ensure the quality of the observed elevation changes, but the cost often limits the amount of in situ data which can be collected. as more tourists are accessing the ice sheets, citizen science might provide in situ data in an environmentally friendly and cost-efficient way. here, we investigate the opportunistic kinematic global positioning system (gps) profiles across the greenland ice sheet, collected by the american-icelandic expedition on the greenlandic icecap 2018. the collected gps data are in good agreement with the widely used nasa’s operation icebridge airborne lidar data measured within ± 10 days, with an average difference of 10.7 cm ± 11.7 cm. the main difference is attributed to changes in the compaction of the snow while driving and changes in the tires’ pressure. the kinematic gps data are then compared with data from the european space agency’s cryosat-2 mission. here, an average bias of 92.3 cm ± 65.7 cm in the two records is observed between the spring cryosat-2 and the truck gps data obtained largely in the dry-snow zone. this suggests that the surface penetration of ku-band radar on the greenland ice sheet and the observed magnitude are consistent with the literature. finally, we compared the 2018 gps data to a profile obtained in 2005 near kangerlussuaq, west greenland. here, the records show an average ice-elevation decrease of 9 m, with peaks at 26 m. these results show that the citizen science kinematic gps data can provide high-resolution data necessary for the validation of satellite altimetry, with the added benefit of potential direct sampling properties of the surface and firn. linking up with citizen-science expeditions is a beneficial way of providing cost-effective satellite validations and may also have a societal impact by involving more people in the climate monitoring of ice sheets. 1 https://doi.org/10.34194/geusb.v47.5369 https://orcid.org/0000-0002-9423-8866 http://orcid.org/0000-0001-8159-6499 https://orcid.org/0000-0002-7288-9545 https://orcid.org/0000-0001-9569-1294 mailto:ssim@space.dtu.dk stokholm et al. 2021: geus bulletin 47. 5369. https://doi.org/10.34194/geusb.v47.5369 2 of 8 www.geusbul let in.org more valuable as societal interest in climate change has increased, including the rapid changes of the marginal zones of the greenland ice sheet (khan et al. 2015). with the growing interest in ice-sheet responses to climate change, validation efforts have become increasingly in demand, and we need to explore new ways to acquire validation data. in the past, the preferred source of validation data for satellite altimetry has been airborne lidar scanning (als). these als measurements started in greenland with the airborne topographic mapper (atm) instrument in 1993; then, they were a part of the nasa’s program for arctic regional climate assessment (parca) programme (krabill et al. 2000) and continued under operation icebridge from 2009 (studinger et al. 2010) to 2019. in addition, the national survey and cadastre (later dtu space) has operated airborne laser scanners in greenland since 2000. this was done in support of, for example, the programme for monitoring of the greenland ice sheet (promice) in 2007, 2011 and 2015 (ahlstrøm et al. 2008; fausto et al. 2012); airborne gravity missions from 2001 to 2003; and numerous esa cryosat-2 validation experiment (cryovex) campaigns since 2003 (skourup et al. 2012, 2013). as airborne validation efforts are costly, here we present a cost-efficient alternative for the purpose of satellite validation, whereby citizen scientists collect kinematic global positioning system (gps) profiles from opportunistic ground traverses on the ice sheet. the icelandic company arctic trucks has demonstrated a unique ability to drive long traverses in antarctica, especially in east antarctica, but also crossing the antarctic plateau from the union glacier to the amundsen coast in west antarctica. in greenland, the ground covered is more limited, with shorter traverses in southern greenland in the 2005, mainly to support the temporary volkswagen winter-test site on the ice sheet near kangerlussuaq. returning to greenland in 2018, arctic trucks completed their first north-south traverse of the greenland ice sheet during the american-icelandic expedition on the greenlandic icecap 2018. the expedition used three modified fourwheel drive trucks (fig. 1c) and seven crew members to demonstrate their capabilities in greenland. the traverse started near isortoq on 19 april 2018 in southern greenland and reached land in northernmost greenland on 29 april 2018 at wulff land. then, it returned south before again exiting the greenland ice sheet on 7 may 2018 near kangerlussuaq, from which the trucks drove over an additional 180 km of ice-free terrain to ship the trucks back to iceland from the coastal town of sisimiut. at the end of the 19-day traverse (8 may 2018), more than 5000 km of inaccessible terrain had been driven (fig. 1a). we used this opportunity to supply the expedition with a geodetic gps receiver in hopes of collecting the valuable ground-truth data of ice-sheet elevation along the track. here, we evaluated the accuracy of the gathered gps measurements and assessed their suitability for providing cost-efficient data to validate satellite altimetry. firstly, we assessed the acquired data in relation to the traditional airborne als data gathered by operation icebridge. after this, we performed an inter-comparison to cryosat-2 satellite data. finally, we showed the long-term capabilities of geodetic gps measurements, by assessing the elevation change near kangerlussuaq between the first measurements in 2005 (i.e. the iceroad route; fig. 1d) and the 2018 arctic truck traverse. 2. data and methods 2.1 arctic trucks kinematic gps measurements a portable geodetic javad tre-3n delta global navigation satellite system (gnss) receiver was mounted on top of one of the arctic trucks vehicles (fig. 1c). the gps continuously logged the position and elevation of the truck at an interval of 2 sec. all gps data were processed with novatel’s waypoint product group post-processing product grafnav in precise point positioning (ppp) mode. this method of gnss processing uses precise orbit and clock information to provide the state-of-the-art positioning for remote locations without base stations or online corrections (from, e.g., real-time kinematic [rtk] services), with a typical accuracy of 5 cm for airborne applications, as demonstrated by, for example, the dtu space cryovex campaigns (skourup et al. 2012, 2013). to link the gps antenna position to the snow surface, the arctic trucks crew repeatedly measured the distance from the snow surface to the antenna reference level, following the tire pressure changes. the values ranged between 227 cm and 243 cm and have been corrected for. the measurements were taken while the truck was stationary, giving rise to further submergence into the snow compared to when the truck was being driven. consequently, kinematic measurements will have a small positive error. additional uncertainties include the varying weight (e.g. consumables such as fuel) of the vehicle further changing submergence, uneven surfaces and the suspension while driving. however, none of these uncertainties were investigated further. unfortunately, no gps data were collected for 6 days during the crossing from wulff land south (31 april–5 may), and the gps data are only available from 19 to 30 april and again from 6 to 8 may 2018. this results in only limited crossover points for the internal validation of the data collected by the gps receiver. the crossover location is seen in fig. 1a, and here, the performance of the geodetic gps is validated to check for significant snow deviations and/or faulty equipment. the crossover samples occur at very close proximity, with many measurements within metres of the ascending track on 21 and 22 april and the descending track on 6 may, which allow for a precise gps height comparison. using the nearest https://doi.org/10.34194/geusb.v47.5369 http://www.geusbulletin.org stokholm et al. 2021: geus bulletin 47. 5369. https://doi.org/10.34194/geusb.v47.5369 3 of 8 www.geusbul let in.org neighbour algorithm with a relatively short euclidean distance of 10 m generates a height difference average of −2.2 cm with a root mean square error (rmse) value of 2.6 cm by comparing 844 crossover points (table 1). consequently, it is assumed that there has been no significant snow elevation change over the survey period and that the gps measurements are accurate to a few centimetres rmse. the spring period of april and may was chosen for the traverse as it offers relatively warm weather, sunlight and stable ice bridges over the crevasses, which are to be crossed when entering and exiting the greenland ice sheet. 2.2 operation icebridge atm data operation icebridge was a nasa programme (2009–2019) that collected airborne remote sensing measurements and incorporated multiple instruments to map ice-surface topography. here, we apply data from the atm instrument package, which includes als, gps and inertial navigation systems. the atm-l2-data (studinger 2018) are a sequence of laser footprints with a size of c. 1 m measured in a circular motion in a swath ranging from 400 m to 1200 m depending on the aircraft height (typically 500–1500 m) and the incidence angle. the nasa atm level 2 elevation, slope, and roughness software condenses the atm point measurements by fitting a plane to blocks of points at regular intervals with 50% overlap. along-track spacing is typically 30 m but is dependent on the aircraft speed. for each alongtrack position, 3–5 evenly spaced across-track points are sampled for either the t2 or t3 atm scanner. each point typically covers 133 m across track when the swath is wulff land (april 30 2018) l d arctic truck crossover sisimiut (may 8 2018) isortoq (april 19 2018) kangerlussuaq nanoq site gps receiver distance to the snow surface a b c d fig. 1d data gap april to may 5 2018 icebridge cryosat-2 lrm cryosat-2 sin arctic trucks iceroad h ei gh t [ m ] 500 1000 1500 2000 2500 3000 fig. 1 map of the expedition route including places visited and dates. a: height data from the geodetic gps mounted on the arctic trucks. b: operation icebridge and cryosat-2 data coverage at elevations above 2000 m across greenland. cryosat-2 data are only shown between 30w and 50w. c: the modified four-wheel drive toyota vehicles used on the arctic trucks expedition. the gps receiver is circled in red (picture courtesy of emil grimsson, arctic trucks, iceland). d: comparison of the 2018 arctic trucks and 2005 iceroad routes. images reproduced from the gebco world map 2014 (www.gebco.net.) https://doi.org/10.34194/geusb.v47.5369 http://www.geusbulletin.org http://www.gebco.net stokholm et al. 2021: geus bulletin 47. 5369. https://doi.org/10.34194/geusb.v47.5369 4 of 8 www.geusbul let in.org 400 m across and 45 m along track. an additional block is measured at the nadir with an across-track length of 80 m and a similar along-track width for a total of 4 or 6 blocks. the vertical accuracy and precision of the atm instrument package are 6.6 cm and 3 cm, respectively (martin et al. 2012). for this study, we selected atm data coinciding within 10 days of the 2018 arctic trucks traverse gps observation (fig. 1b) measured at altitudes above 2000 m where the ice sheet is most level. this ensures limited weather-induced changes in the snowpack between the two types of observations and minimises the potential effects of the surface slope. crossovers between operation icebridge and arctic trucks are found by applying a nearest neighbour algorithm, locating all operation icebridge values within a selected euclidean distance of the arctic trucks points. to limit the potential elevation deviation from terrain slopes, both operation icebridge and arctic trucks data are subtracted from a digital elevation model (dem) using the greenland ice mapping project (gimp) dem (howat et al. 2014) and bilinear interpolation. the gimp dem from 2007 has a spatial resolution of 30 m with an error ranging from ± 1 m on the greenland ice sheet to ± 30 m in areas with high relief (howat et al. 2014). the final crossover values are found as the average of these neighbours for each point: h n h demoib nn n n nnn = − = ∑1 1 . where hn is the elevation of the operation icebridge neighbours, hoib is the crossover elevation of operation icebridge and nnn is the numbers of nearby neighbours. we define the measurement accuracy as: ∆h h dem hoib at oib= −( ) − ( ) where hat is the elevation of the arctic trucks measurements with the dem subtracted. the operation icebridge measurements are in general located further away from the arctic trucks track. to ensure sufficient measurements, a higher euclidean distance for the nearest neighbour algorithm is necessary. in fact, the error, δhoib, scales as a fermi dirac-like function of the distance, and using the relative optimum of 75 m ensures that most data have a minimal δhoib. increasing the distance will enlarge δhoib as more points further away from the track are included, introducing surface elevation changes. we then defined the arctic trucks measurement accuracy (εat) as the average of all δhoib. εat, which incorporates the uncertainties associated with driving, local terrain variations and gps inaccuracies. 2.3 cryosat-2 – satellite data esa’s cryosat-2 satellite was launched on 8 april 2010 into a full repeat of its orbit in 369 days and with a sub-cycle of 30 days. the main instrument onboard the satellite is the synthetic aperture radar (sar)/interferometric radar altimeter (siral). for the area of greenland, the siral instrument operates in two modes: (1) in the interior parts of the ice sheets, it operates in the low-resolution mode (lrm), which is equivalent to the conventional beam-limited mode applied by previous ku-band radar altimeters; (2) over the coastal areas, siral is switched into the sar interferometric (sarin or sin) mode, in which the dual antennas provide a directionally determined radar return across-track and a high-resolution synthetic aperture processing along-track (bouzinac 2012). we use both lrm and sin mode as processed in the official esa level 2 processing-chain baseline cacs team & mssl team. (2015). the pulse-limited footprint of cryosat-2 in lrm mode is about 2.15 km2, with a width of 1.65 km (bouzinac 2012). further, we limit the cryosat-2 data from 9 april to 18 may 2018 (fig. 1b), corresponding to the timing of the arctic trucks traverse of ± 10 days. similar to operation icebridge, the data are limited to elevations above 2000 m. with the estimate of εat from the operation icebridge analysis, arctic trucks gps data can be compared to cryosat-2, and crossovers are found in a similar fashion table 1 statistics for the arctic trucks kinematic gps data (572 607 observations) compared with operation icebridge, cryssat-2 and iceroad data. n mean [m] med. [m] std [m] max [m] min [m] rmse [m] d [m] gps validation 844 -0.022 -0.021 0.015 0.057 -0.072 0.026 10 δhoib 1217 -0.107 -0.110 0.117 0.280 -0.648 0.159 75 δhcs2 10 463 0.923 0.905 0.657 5.619 -5.631 6.764 75 δhiredge 104 -16.995 -16.385 3.842 -9.644 -26.042 16.158 75 δhirinterior 277 -9.035 -9.713 6.940 9.900 -25.929 11.392 n denotes the number of points in the derived statistics (mean, med.: median; std.: standard deviation). rmse: root mean square error. d denotes the distance used in the nearest neighbour algorithm. cryosat-2 outliers larger than 3 standard deviations have been removed. oib: operation icebridge; cs2: cryosat-2; ir: iceroad. https://doi.org/10.34194/geusb.v47.5369 http://www.geusbulletin.org stokholm et al. 2021: geus bulletin 47. 5369. https://doi.org/10.34194/geusb.v47.5369 5 of 8 www.geusbul let in.org to the operation icebridge study, by subtracting the 2007 gimp dem and applying the nearest neighbour algorithm with a euclidean distance of 75 m. using a lower euclidean distance, for example, 50 m, decreases the amount of available crossover points by 94%. afterward, cryosat-2 (cs2) is validated by calculating the height difference: ∆h h dem hcs at cs2 2= −( ) − ( ) 2.4 the 2005 iceroad data the iceroad data were collected in may 2005 on another opportunity truck traverse, as the volkswagen winter-test site on the ice sheet was closing down. the gps profile was measured along the iceroad from point 66 at the ice-sheet margin near kangerlussuaq to the volkswagen winter-test site location, called nanoq (fig. 1d). an ashtech z-12 mobile receiver with a sampling interval of 5 sec was used and mounted on top of the vehicle, similar to arctic trucks, with a measured height of 2.45 m, which has been corrected for. the compacted ice road was slightly elevated in comparison to the surrounding snow surface. the gps data have been post-processed using precise orbit and clock information (waypoint software in ppp mode). the formal uncertainty estimate from this processing is 0.3 cm, but the absolute vertical accuracy of gps solutions in kinematic mode is often found to be around 5–10 cm under arctic conditions (sørensen et al. 2018b). after reprocessing, it became obvious that the 2005 and 2018 traverses intersect for about 300 km, in the ablation zone of the greenland ice sheet near kangerlussuaq (fig. 1d). here, we compared the 2005 iceroad and 2018 arctic trucks traverses in two steps. the overlapping region illustrated in fig. 1d is found using the nearest neighbour algorithm with a distance of 75 m, which seems to be a good fit between the spatial proximity and elevation difference. further inland, the two truck routes diverge and are separated by kilometres in latitude, and therefore, the crossovers method is not appropriate. instead, both tracks are compressed into blocks based on the overlapping longitude as the traverse is largely from east to west. each block incorporates all points with a 100 m interval in longitude yielding 277 blocks where the longitude, latitude and elevation have been averaged into one point. afterward, the slope-dependent elevation is removed by correcting the dem differences for both the iceroad (ir, in equation 4) and the arctic trucks data with the gimp dem. thereby, the elevation changes from 2005 to 2018 are estimated for the interior portion of the overlapping tracks: ∆h h dem h demir ir atinterior = −( ) − −( ) 3 results the distribution of all validation targets is shown in fig.  2, and the accompanying statistics are listed in table  1, in which we include the number of samples, mean, median, standard variance, max., min., rmse and the euclidean distance in the nearest neighbour algorithm. these statistics are based on the data illustrated in fig. 1. from the statistics of the different inter-comparisons given in table 1, it is evident that δhoib has fewer data points than cryosat-2 despite having a much higher sampling frequency. this is partly because most of the operation icebridge data are near the coast (typical for their acquisitions and flying tracks), leaving few overlapping areas with arctic trucks, while cryosat-2 has a larger coverage. we find that the elevations measured with operation icebridge and arctic trucks are mostly similar, with relatively small outliers, while in comparison, cryosat-2 fig. 2 histograms of the elevation differences (h) between the arctic truck acquisition and data from a: operation icebridge hoib has a bin size of 1 cm and there are no significant data clusters outside the range -0.7 to 0.3 m, b: cryosat-2 hcs2 has a 2 cm bin size, with no significant clusters outside -0.5 to 3.25 m, c: iceroad hir in the range of -26 to 12 m with a bin size of 1 m; no data occur outside of this range. ice edge inland ∆h [m] ∆h [m]∆h [m] n um be r o f s am pl es 0 20 40 60 80 100 –0.6 –0.4 –0.2 0.0 0.2 –0.5 0 0.5 1.0 1.5 2.0 2.5 3.0 0 200 400 600 800 1000 1200 1400 0 5 10 15 20 25 –25 –20 –15 –10 –5 0 5 10 a b c https://doi.org/10.34194/geusb.v47.5369 http://www.geusbulletin.org stokholm et al. 2021: geus bulletin 47. 5369. https://doi.org/10.34194/geusb.v47.5369 6 of 8 www.geusbul let in.org has large outliers, despite removing outliers ± 3 standard deviations from δhcs2. similar means and medians indicate that most overlapping data points are relatively close to the δh for both operation icebridge and cryosat-2. based on the average δh, we see that operation icebridge generally measures a higher elevation and cryosat-2 measures a lower elevation compared to arctic trucks. the average of δhoib is 10.7 cm, and the histogram in fig. 2a illustrates that nearly all data points are located in the range of –30 to 10 cm, with a wide and tall cluster. the mean and median of operation icebridge are very similar, indicating that the mean is not significantly affected by large outliers. the cryosat-2 histogram contains a main cluster from 60 to 105 cm region with two small clusters at 175–190 cm and at 290–305 cm. if the large outliers were not removed from the cryosat-2 data, the mean would be significantly altered. the elevation differences for the iceroad data (fig. 2c) show a different distribution for measurements near the ice sheet edge compared to the interior measurements. the spread at the ice edge is attributed to the non-uniform elevation change in this part of the ice sheet, whereas the larger spread in the interior observations can be attributed to the divergence of the two traverse routes. discussion operation icebridge atm data have been used to assess the performance of the arctic trucks gps observations of the ice-sheet surface elevation, and the results show an average difference of 10.7 cm ± 11.7 cm. the bias is likely from a combination of arctic trucks uncertainties, the uncertainties from the atm data and the footprint coverage of the atm icessn point compression. regarding the arctic trucks measurements, we note that the main uncertainties are the tendency of the wheels to compress the uppermost snow layer, the uneven surfaces and some counteraction of the vehicle suspension system while driving. the tire pressure is generally low to provide a larger surface grip, and it is adapted to snow conditions as necessary (low for soft snow and high for solid surfaces). regarding the submergence into the snow, we argue that the mean error is not zero, because the height measurement of the elevated gps location on the truck is taken while in a static position, where the truck will submerge the most. once the truck enters a kinematic state, the compaction will often decrease, caused by the wheels moving onto fresh snow, which is replaced repeatedly. when following the tracks of the other trucks, the snow will be compressed multiple times, which can lead to greater snow compaction. additionally, the measured distance from the snow surface to the gps receiver could vary in accuracy and depends largely on how frequently the gps position is measured. the tall and wide cluster of δhoib in fig. 2a has no spatial geographical dependency but is randomly distributed over the ice sheet. locally varying snow conditions could explain the overlapping distributions, as the uncertainty of the distance measured from the snow surface to the gps receiver and the submergence of the wheels affect the readings. if the distance from the snow surface to the gps is measured while on hard, sturdy snow and the vehicle later enters soft snow, the submergence gives rise to a negative error contribution. therefore, the tall cluster could simply be a result of the local snow conditions being similar to the point where the distance from the snow surface to the receiver was measured. this would imply that the wide cluster stems from varying local snow conditions affecting the measured distance, which also explains why this cluster is much wider than the tall cluster. the vertical accuracy of the lidar in the atm instrument package is estimated to be 6.6 cm (martin et al. 2012). factoring in this uncertainty and comparing it to the δhoib of 10.7 cm, it is evident that distinguishing measurement uncertainties from small elevation deviations in the snow is difficult. also, each operation icebridge measurement point covers 133 m × 45 m, and with a worst-case nearest neighbour coverage of 283 m × 193 m, using a distance of 75 m would certainly allow for additional errors in the range of centimetres. however, this is mitigated by weighting the neighbour distance and atm level 2 point compression. a larger quantity of data from operation icebridge across the greenland ice sheet could lower the nearest neighbour range and possibly improve compliance between the two measuring approaches. based on this discussion, we estimate εat to be 10.7 cm. comparing the arctic trucks and cryosat-2 data results in an average δhcs2 of 92.3 cm ± 65.7 cm, as seen in fig. 2b. we observed 76 large outliers outside the ± 3 standard deviations and these were removed. the two small clusters to the right of the main distribution are both from two independent spatial geographical locations. the mean of δhcs2 deviates by a factor of 10 compared to δhoib, which is significantly larger than the estimated εat and therefore another aspect that clearly influences the result. it has been theorised that the ku-band of the cryosat-2 siral radar altimeter has a penetration depth of 5–12 m for dry snow snow (nilsson et al. 2015; remy et al. 2015). here, we use the penetration depth to describe the height bias between the radar altimeter and the surface. significant snow penetration was observed in 2012, when a greenland-wide ice-melt event increased the apparent elevation of large areas of the ice sheet, with up to 2 m differences from june to august, which indicated that the snow penetration depth had significantly decreased. this was caused by the formation of https://doi.org/10.34194/geusb.v47.5369 http://www.geusbulletin.org stokholm et al. 2021: geus bulletin 47. 5369. https://doi.org/10.34194/geusb.v47.5369 7 of 8 www.geusbul let in.org refrozen melt layers, which raised the reflective surface (nilsson et al. 2015) and scattered the sar wave closer to the actual surface. however, this is also highly dependent on the type of processing algorithm and retracker (sørensen et al. 2018b). another estimate of the dry snow penetration assessed it to be 1 m; it was a magnitude lower for wet snow (ulaby & stiles 1981). the snow penetration explains the compared positive deviations, as we would expect the cryosat-2 data to be lower than the arctic trucks data. there are, however, 176 crossover instances where the difference is negative. since this large divergence only occurs with the cryosat-2 data and without spatial correlation, we must conclude that this positive bias stems from either faulty measurements in these data or the esa level-2 processing-chain baseline c. we experience similar deviations, where cryosat-2 values are much higher than the als, in a study by sørensen et al. (2018b). using a different processing chain can limit the surface penetration of the radar (slater et al. 2019); hence, applying a dedicated retracker could improve the coherence between arctic trucks and cryosat-2, by moving the cryosat-2 observations closer to the surface. despite the uncertainty in the processing chain, our analysis agrees with the results of sørensen et al. (2018b) for the same processing chain, as the mean and median observed data are very similar for a flat area. the exact snow penetration is difficult to estimate, since temperature fluctuates significantly across the ice sheet depending on elevation, weather and latitude. therefore, an assessment for the average snow penetration of cryosat-2 for this april – may time period is about 90 cm. finally, we investigate δhir. the two data acquisitions are very similar in setup, except that the 2005 iceroad traverse drove along an existing groomed snow road with fewer uncertainties, for example, no submergence. the road was slightly elevated, but this was considered when processing the iceroad gps data. from the δhir statistics (table 1), it is apparent that a significant amount of ice has vanished from 2005 to 2018, particularly close to the edge of the ice sheet. there has been a mean loss at the edge of 17 m and up to 26 m. it is evident from fig. 2c that the loss is more consistent closer to the ice-sheet edge. there are also multiple instances of increased height. further inland, the mean loss is 9.04 m, which is similar to other estimates of the surface elevation change in this area, for example, the esa’s climate change initiative (cci) surface elevation change product (simonsen & sørensen 2017; sørensen et al. 2018a). investigating the arctic trucks and iceroad deviations from the gimp dem reveals an increasing difference from the volkswagen test site toward the edge of the greenland ice sheet. this indicates a steepening of the ice sheet moving westward, implying higher melt rates at the margins of the ice sheet than inland, which is increasingly observed across the greenland ice sheet (sørensen et al. 2018a). 5 conclusions here, we have shown how opportunistic data gathered by citizen scientists can provide valuable validation data to improve satellite altimetry measurements. the data have been found to be within 10.7 cm ± 11.7 cm of the traditional and more costly als data provided by the nasa’s operation icebridge. this error estimate formed the basis for using the data for satellite validation. here, coincident cryosat-2 observations, from the esa level-2 baseline c processing chain, showed a mean snow penetration at the ku-band of 92.3 cm ± 65.7 cm, which agrees with other studies applying traditional airborne als. finally, the long-term durability of this opportunistic citizen science was proven by comparison to a similar truck traverse in 2005 around kangerlussuaq. the results showed an average surface elevation change of 17 m and up to 26 m near the edge of the ice sheet. further inland, the mean change is 9.04 m, and we observed an increase of the slope toward the ice-sheet edge. with a precision similar to operation icebridge, gps profiles from vehicle traverses are a strong contender for collecting independent data to validate satellite and airborne height measurements, with the additional benefit option of providing other in situ data from, for example, snow pits and firn. therefore, if geodetic gps receivers can be supplied to opportunistic citizen scientists, we see this as a cost-efficient and valuable source of data to validate altimetry satellites. such initiatives allow citizen scientists to compare data from one expedition to the next and see first-hand the large changes happening on the ice sheet. in this way, it could further help to engage the public and raise awareness of the large changes underway in greenland. acknowledgements we thank all contributors and participants in the american-icelandic expedition on the greenlandic icecap 2018 for giving us the opportunity to study the use of kinematic gps as a validation tool. we thank emil grimsson, arctic trucks, for allowing the mount of the dtu gps receiver on one of the expedition trucks and for the efficient installation and field operations during the traverse by the arctic trucks team. the operation icebridge data were downloaded from the national snow and ice data center, and the esa cryosat data were downloaded from the esa data hub. iceroad 2005 data were collected by a dtu team of rene forsberg and henriette skourup, also on an opportunity basis, courtesy of the nanoq snow-road maintenance team. author contributions rf, smh and sbs conceived the idea and outlined the study. as and sbs conceived the methodology and conducted the analysis, data processing and writing of the original draft. smh and rf prepared the gps equipment for the field deployment. all authors contributed to the final writing, reviewing and editing. https://doi.org/10.34194/geusb.v47.5369 http://www.geusbulletin.org stokholm et al. 2021: geus bulletin 47. 5369. https://doi.org/10.34194/geusb.v47.5369 8 of 8 www.geusbul let in.org references acs team & mssl team. 2015: cryosat ground segment, instrument processing facility – l2 products format specification (l2-fmt). baseline c version. advanced computer systems, doc. no.: cs-rs-acs-gs-5123, 98 pp. https://earth.esa.int/web/guest/missions/cryosat/ipf-baseline/-/article/cryosat-ground-segment-instrument-processing-facility-l2-l2-products-format-specification-l2-fmt-6636 ahlstrøm, a.p. et al. 2008: a new programme for monitoring the mass loss of the greenland ice sheet. geological survey of denmark and greenland bulletin 15, 61–64. https://doi.org/10.34194/geusb.v15.5045 bouzinac, c. 2012: cryosat product handbook. report no., esrin esa and mullard space science laboratory – university college london, esa esrin, frascati, italy. fausto, r.s., van as, d. & promice project team. 2012: ablation observations for 2008–2011 from the programme for monitoring of the greenland ice sheet, promice. geological survey of denmark and greenland bulletin 26, 73–76. https://doi.org/10.34194/geusb. v26.4765 forsberg, r., sørensen, l. & simonsen, s. 2017: greenland and antarctica ice sheet mass changes and effects on global sea level. in: cazenave, a. et al. 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data and methods 2.1 arctic trucks kinematic gps measurements 2.2 operation icebridge atm data 2.3 cryosat-2 satellite data 2.4 the 2005 iceroad data 3 results discussion 5 conclusions acknowledgements author contributions references table table 1 statistics for the arctic trucks kinematic gps data (572 607 observations) compared with op figure fig. 1 map of the expedition route including places visited and dates and equipment. see the figure fig. 2 histograms of the elevation differences between the arctic truck acquisition and satellite d sandrin et al. 2020: geus bulletin 44. 5240. https://doi.org/10.34194/geusb.v44.5240 1 of 6 research article | short geophysics for urban mining and the first surveys in denmark: rationale, field activity and preliminary results alessandro sandrin1, aleksandar maricak2, björn h. heincke1, rune j. clausen1, lars nielsen2, jakob k. keiding1 1centre for minerals and materials (mima), geological survey of denmark and greenland, copenhagen, denmark. 2department of geosciences and natural resource management, university of copenhagen, copenhagen, denmark abstract geophysical methods have been widely used in recent decades to investigate and monitor landfill sites for environmental purposes. with the advent of the circular economy, waste contained in old landfills may be considered a resource that can be developed. since the content of old landfills is largely unknown, the occurrence and quantity of valuable materials must be investigated before embarking on any development activity. two landfills on sjælland, denmark (located at hvalsø and avedøre) were selected for a pilot study to characterise their content. at both locations, a set of geophysical surveys is underway. here, we present the data obtained from magnetic and 2d seismic refraction surveys. magnetic data show various anomalies that can be interpreted as caused by iron-rich waste. at both sites, the landfill material results in generally low p-wave velocity (<400 m/s), lower than those obtained for quaternary sediments at avedøre. the seismic velocities appear to increase in the presence of metals or by compaction with depth (>550 m/s). we propose that seismic refraction can thus define the bottom of the landfill and possibly its internal structure, especially when combined with other methods. introduction societies around the world are increasingly looking at renewable energy and recycling of materials in response to a changing global economy and environmental challenges. within the field of ‘circular economy’, landfill mining has gained momentum in recent years (e.g. wagner & raymond 2015, kieckhäfer et al. 2017, puthussery et al. 2017). recycling materials from existing landfills may seem straightforward; nevertheless, the economic value of commodities hidden in the waste should be carefully evaluated before they can be exploited. to determine the occurrence and potential quantity of recyclable materials in a landfill, non-invasive methods are preferred (green et al. 1999; cardarelli & di filippo 2004; balia & littarru 2010; boudreault et  al. 2010; belghazal et al. 2013; dumont et al. 2017; di maio et al. 2018). digging or drilling into landfills may be expensive and environmentally hazardous. for this reason, we chose geophysical methods to investigate two selected landfills on sjælland, denmark (located at hvalsø and avedøre; fig. 1). to our knowledge, these are the first such surveys of landfills in denmark. in this article, we briefly explain the rationale behind the project and its contribution towards *correspondence: alsa@geus.dk received: 23 jan 2020 accepted: 15 apr 2020 published: 02 july 2020 keywords: geophysics, urban mining, circular economy, landfill, denmark abbreviations: rms: root mean square sirt: simultaneous iterative reconstruction technique geus bulletin is an open access, peerreviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: catherine jex (geus) reviewed by: david caterina (université de liège, belgium), philip carpenter (northern illinois university, usa) author contributions: see page 6 funding statement: see page 6 competing interests: none declared  additional files: none provided https://doi.org/10.34194/geusb.v44.5240 https://creativecommons.org/licenses/by/4.0/deed.ast sandrin et al. 2020: geus bulletin 44. 5240. https://doi.org/10.34194/geusb.v44.5240 2 of 6 www.geusbul let in.org a sustainable economy. we describe the field activities at both sites and present the first results from magnetic and 2d seismic refraction surveys of the landfills. rationale behind landfill mining raw materials are an essential part of our modern society and there is a growing demand for mineral raw materials in europe and the rest of the world. this increasing demand raises growing concerns regarding the availability of these resources. in addition to supply risk and resource depletion, our large consumption of resources results in pollution and excessive land use, increasingly jeopardising earth’s life-support systems (jackson 2009; wwf 2014). to address these issues, and related sustainability needs, the circular economy has become an increasingly popular concept, which aims to eliminate waste and continual use of resources. landfills have historically been considered a practical and cost-efficient method for final storage of waste. however, landfills represent a potential environmental hazard and suboptimal use of resources. the eu is estimated to contain up to 500  000 landfills (jones et al. 2013) and although measures have been taken to increase recycling and waste disposal by incineration, 40% of waste in the eu-27 (the 27 member states of the eu) is still landfilled (blumenthal 2011). therefore, landfills continue to be a significant leakage in the circular economy. landfill mining refers to the excavation, processing, treatment and recovery of deposited materials situated in informal waste dumps or in structured landfills (savage et al. 1993). the concept was introduced in the early 1990s and in most cases, it was limited to extraction of methane or for land reclamation. landfill mining is now being further developed with a greater focus on resource recovery of raw materials. at first glance, landfill mining seems an ideal solution that combines remediation,  land reclamation and extraction of raw materials, fig. 1 landfill sites selected for this study. a: hvalsø (1 shows the area investigated), b: avedøre (2 shows the area investigated; b marks the location of an existing borehole), c:  location of the two landfills in denmark (digital elevation shown in grey shading). the landfill in hvalsø is composed of various kinds of household waste including metallic objects. the area was remediated and covered with c. 2 m of soil. in avedøre, the waste is at the surface and very heterogeneous in character, consisting of wood, plastic, metal and rubber. https://doi.org/10.34194/geusb.v44.5240 http://www.geusbulletin.org sandrin et al. 2020: geus bulletin 44. 5240. https://doi.org/10.34194/geusb.v44.5240 3 of 6 www.geusbul let in.org but  it has not been broadly implemented. one of the main challenges is the often-low economic performance of landfill mines (e.g. frändegård et al. 2015; laner et al. 2019; clausen & kalvig 2020). prospecting for the most profitable landfills is an important incentive to recognise the feasibility of landfill mining (krook et al. 2012). field work and methods the first site is an old landfill in hvalsø, sjælland, denmark (55°35́ 14́ ń, 11°52́ 16́ é; fig. 1). the waste was dumped in an abandoned gravel pit between 1980 and 1985 (roskilde amt teknisk forvaltning 2003). the site was remediated and covered with soil and a number of wells currently produce biogenic gas from within the landfill (biorem 2016). documentation on the quantity and quality of the various types of waste is limited. however, personal and anecdotal information indicate the likely presence of metals in some parts of the landfill. their exact location and amount are unknown. the second site is located in avedøre, sw of copenhagen, sjælland, denmark (55°36́ 17́ ń, 12°28́ 13́ é; fig. 1) and is an active landfill. the waste has not been buried, and is therefore still observable at the surface. the waste extends to c. 5 m depth, with a 20 cm gravel layer at the base for drainage. below the gravel, c. 10–13 m of glacial soil lies directly above carbonates of the chalk group (borerapport dgu 208. 3888, geus 1999). the waste is shredder residue of household origin with variable content of plastic, metals, wood and rubber (av miljø, personal communication 2019). magnetic surveys were conducted to define areas with high magnetic anomalies, likely related to the presence of fe-bearing waste (marchetti et  al. 2013). the magnetic data were acquired using a geometrics g-858 magmapper instrument, with two caesium vapour sensors. a differential gps positioning system was attached to the operator, who walked the entire area in straight lines whenever possible. magnetic data were processed, reduced-to-the-pole and mapped using geosoft® software. we conducted 2d seismic refraction surveys to delineate the depth to fe-bearing waste at both sites, assuming that p-wave velocities would increase in fe-bearing waste compared to the surrounding waste (e.g. lanz et al. 1998; de iaco et al. 2003). surveys were conducted along a 115-m long transect. a single geode seismic recorder with 24 channels was used and geophones were deployed with 5 m spacing. the seismic source for the survey was a hammer hitting a plate coupled to the ground (four stacks). the distance between shots was 2.5 m at hvalsø, and 5 m at avedøre. for each shot, pre-processing of the data, manual picking of the first-arrival travel times and tomographic inversion were performed using the software package reflexw® by sandmaier geophysical research (sandmeier 2014). two-dimensional first-arrival travel-time tomography is based on an iterative non-linear deterministic inversion and uses rectangular cells for both forward modelling and inversion (sandmeier 2014). in the forward modelling, a finite-difference approximation of the eikonal equation is used to calculate the travel-time field for the ray tracing. determination of model updates relies on an iterative sirt algorithm (simultaneous iterative reconstruction technique; sandmeier 2014). preliminary results and discussion magnetic surveys magnetic data in both sites show great variability (47  000–56  000 nt), indicating that fe-content is dispersed heterogeneously throughout the waste. at hvalsø, a significant local magnetic anomaly was detected in the western part of the landfill (white arrow in fig. 2a). the strong anomalies detected in the southern part of the site are likely produced by nearby manmade structures (e.g. high voltage lines and pylons, fig. 2a). at avedøre, the magnetic anomalies do not seem to indicate such local enrichment in fe-bearing materials. rather a more uniform distribution is observed, with higher values (up to 55 000 nt) at the east and lower values (down to 48 000 nt) at the west of the site (fig. 2b). lower total magnetic field values are observed outside the landfill. 2d seismic refraction surveys the first-arrival travel times were used for the modelling of the data (figs 3a, b, c). a starting model with a velocity gradient of 350 m/s at the surface, continuously increasing with depth, was created (fig. 3d). the final tomographic model is shown in fig. 3e. ray-tracing diagrams calculated for the given velocity field (fig. 3f) show the sub-surface coverage provided by the first-arrival travel times. for the seismic refraction data recorded at avedøre, the same approach was used. first, the seismic data for each shot were loaded in reflexw, and then manual picking of the first arrivals was conducted. band-pass filters and various scaling parameters were applied to allow a reliable interpretation of the first breaks (figs 4a, b, c). a simple one-layer 2d starting model was created with a 250 m/s p-wave velocity at the surface and increasing with depth (fig. 4d). various tests were carried out, and in this article only one final model is presented (fig. 4e). the tomographic inversion shows reasonable root mean square (rms) values (c. 12 ms – comparable to the picking uncertainty), especially at far offsets. figure 4f displays the ray coverage obtained https://doi.org/10.34194/geusb.v44.5240 http://www.geusbulletin.org sandrin et al. 2020: geus bulletin 44. 5240. https://doi.org/10.34194/geusb.v44.5240 4 of 6 www.geusbul let in.org and fig. 4g compares the measured and calculated travel times. the velocity field obtained from seismic tomography is in agreement with the lithological log from a nearby borehole (fig. 4e), which records c. 13 m of glacial soil overlying cretaceous carbonates of the chalk group. the seismic velocities of the waste (uppermost 5 m in fig. 4e) are in the range 250–320 m/s. the p-wave velocity increases to >350 m/s at c. 5 m depth (red hatched line in fig. 4e). at c. 18–20 m below the surface, the velocity appears to exceed 800 m/s. seismic velocity distributions from both landfills indicate velocities of 250–400 m/s for non-compacted waste (0–8 m burial depth) of mixed origin. similar velocities were obtained for landfills in switzerland (300–800 m/s; lanz et al. 1998; de iaco et al. 2003) and usa (350–550 m/s; carpenter et al. 2013). konstantaki et  al. (2016) obtained particularly low p-wave velocities (150–200 m/s) in a landfill in the netherlands, and attributed these to the presence of biogas (which could be the case at hvalsø). fig. 2 total magnetic field maps for landfill sites. a: hvalsø, b: avedøre. for reference, the total magnetic field in copenhagen is 50  360 nt. seismic lines are shown in black. fig. 3 two-dimensional seismic refraction data at hvalsø. a, b and c: data from three selected shots. d: starting model. e: final p-wave velocity field (vp) obtained from tomographic inversion. the black hatched line marks the top of the fe-bearing waste. f: ray coverage. g: comparison between calculated (dots) and observed (lines) first-arrival travel times. rms: root mean square. https://doi.org/10.34194/geusb.v44.5240 http://www.geusbulletin.org sandrin et al. 2020: geus bulletin 44. 5240. https://doi.org/10.34194/geusb.v44.5240 5 of 6 www.geusbul let in.org waste-containing metals are expected to show relatively high p-wave velocities. if the surrounding waste has a velocity in the order of 300 m/s and with velocities of metals higher than 3000 m/s (salisbury et  al. 1996; reichmann & jacobsen 2004; nourani et al. 2017), then a mixed volume of the two could have velocities in the order 500–600 m/s – depending upon the amount and connectivity of the metals. at avedøre (fig. 4), velocities of 800–900 m/s in the deeper and central parts of the profile can be attributed to the presence of ‘fast’ chalk lithologies, which often have p-wave velocities >2000 m/s (nielsen et al. 2011). here, the ray coverage at depth is not optimal, thus these velocity estimates have large uncertainties. moreover, the erosion and weathering of the uppermost chalk during glaciations might have led to decreased seismic velocities. considering these two factors, we can only conclude that the chalk at avedøre has p-wave velocities in excess of 800 m/s. conclusions and further work preliminary results for two test sites in denmark indicate waste with p-wave velocities between 250 and 450 m/s – comparable to other lose sediments that have not undergone intense compaction. the presence of metals within the waste, however, increases the seismic velocity to values higher than 500 m/s. the waste appears to be more magnetic than the surrounding glacial soil, with higher values corresponding to areas of  high p-wave velocities. the analysis of seismic data alone can support the definition of the landfill base and  provide useful information on waste properties, whereas the magnetic survey seems promising for delimiting the deposits horizontal extent, mapping internal heterogeneity and identification of iron-rich zones. integration with additional geophysical data is required for a more comprehensive characterisation and more detailed site delimitation. for example, geoelectric data can now be modelled to constrain the  3d inversion of the magnetic data at hvalsø and avedøre. further magnetic susceptibility measurements directly on the waste at avedøre, will also better  constrain the modelling of the magnetic field data here. acknowledgements the authors thank the department of geosciences and natural resource management (ign) at the university of copenhagen for providing the geophysical instruments. peer jørgensen and mia benner from ign are thanked for the help provided in the field. per kalvig from geological survey of denmark and greenland (geus) is thanked for discussions and comments on an earlier version of the manuscript. region sjælland and av miljø are acknowledged for making the sites available and for providing key information on the landfills. suggestions by the two reviewers helped to improve the original manuscript. additional information funding statement the study was funded by centre for minerals and materials (mima) and the geological survey of denmark and greenland (geus). fig. 4 two-dimensional seismic refraction data at avedøre. a, b and c: data from three selected shots. d: starting model. e: final p-wave velocity (vp) field obtained from tomographic inversion. the red hatched line marks the interface between waste and glacial soil, white the hatched line marks the top of chalk. a simplified lithological log of a nearby borehole is shown on the right (light brown shading: glacial sediments, white: carbonates). f: ray coverage. g: comparison between calculated (dots) and observed (lines) first-arrival travel times. rms: root mean square. https://doi.org/10.34194/geusb.v44.5240 http://www.geusbulletin.org sandrin et al. 2020: geus bulletin 44. 5240. https://doi.org/10.34194/geusb.v44.5240 6 of 6 www.geusbul let in.org author contributions as: fieldwork, project planning, data analysis, lead scientist. am: seismic acquisition and processing in hvalsø. bhh: supported data modelling. rjc: supported data acquisition. ln: supported data modelling. jkk: writing of the circular economy section and team leader. references balia, r. & littarru, b. 2010: geophysical 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https://doi.org/10.2138/am-2004-0718 https://doi.org/10.2138/am-2004-0718 https://www.danishsoil.org/media/test_sites/7/documents/historisk%20redeg%c3%b8relse2.pdf https://www.danishsoil.org/media/test_sites/7/documents/historisk%20redeg%c3%b8relse2.pdf https://www.danishsoil.org/media/test_sites/7/documents/historisk%20redeg%c3%b8relse2.pdf https://doi.org/10.2113/gsecongeo.91.5.821 https://www.sandmeier-geo.de/download/reflexw_manual_a4.pdf https://www.sandmeier-geo.de/download/reflexw_manual_a4.pdf https://doi.org/10.1016/j.wasman.2015.06.034 https://www.worldwildlife.org/pages/living-planet-report-2014 research article knudsen et al. 2021: geus bulletin 47. 5336. https://doi.org/10.34194/geusb.v47.5336 1 of 11 fingerprinting sources of salinity in a coastal chalk aquifer in denmark using trace elements christian knudsen* , klaus hinsby , rasmus jakobsen , lars juul kjærgård, per rasmussen geological survey of denmark and greenland, copenhagen, denmark abstract salinity levels above the drinking water standard (>250 mg/l cl–) are observed at shallow depth in a maastrichtian chalk aquifer on the island of falster, south-eastern denmark. to understand the source of the salt, 63  samples from 12 individual, 1 m, screened intervals between 14 and 26 m b.s. were collected from 1 may to 4 june 2018. the samples were collected during a tracer test to estimate the dual porosity properties of the chalk and were analysed for a wide range of elements. furthermore, samples from the baltic sea and from deeper saline aquifers in the area (40 and 85 m b.s.) were analysed for comparison. the geochemical data were analysed using an unsupervised machine-learning algorithm, self organising maps, to fingerprint water sources. the water composition in the screened intervals at various stratigraphic levels has specific geochemical fingerprints that are maintained for the first days of pumping and are distinct amongst the different levels. this suggests an evolution in water composition because of reaction with the chalk. water composition is distinct from both seawater from the nearby baltic sea and salty water from deeper levels of the reservoir. thus, neither up-coning of salty water nor intrusion of seawater caused the elevated salinity levels in the area. the slightly saline composition of groundwater in the shallow aquifer (14–26 m b.s.) is more likely because of incomplete refreshing of the salty connate water in the chalk during the pleistocene and holocene. furthermore, the geochemical fingerprint of salty water from the deeper aquifer at 40 m was similar to water from the baltic sea, suggesting a baltic sea source for salt in the aquifer at 40 m b.s., c. 100 m from the coast. statistical analysis based on self-organising maps is an effective tool for interpreting a large number of variables to understand the compositional variation in an aquifer and a useful alternative to linear dimensionality-reduction methods such as principal component analysis. the approach using the multi-element analysis combined with the analysis of self-organising maps may be useful in future studies of groundwater quality. 1 introduction sustainable water resources management and saltwater intrusion in coastal aquifers are huge challenges globally, and these challenges are projected to be more severe in the future due to climate change and sea-level rise (hinsby et al. 2016; rasmussen et al. 2013). *correspondence: ckn@geus.dk received: 20 aug 2020 accepted: 12 mar 2021 published: 23 july 2021 keywords: fingerprinting, groundwater, chalk aquifer, trace elements, selforganising map abbreviations: b.s.: below surface bmu: best matching units dgu: danish geological survey geus: geological survey of denmark and greenland icp-ms: inductively coupled plasma mass spectrometry som: self-organising map geus bulletin is an open access, peerreviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: julian koch (geus, denmark) reviewed by: søren munch kristiansen (aarhus university, dk), sascha müller (university of copenhagen, dk), boris van breukelen (tu delft, nl) funding: see page 11 competing interests: none declared additional files: see page 11 https://doi.org/10.34194/geusb.v47.5336 https://orcid.org/0000-0003-4722-7678 https://orcid.org/0000-0003-1190-4550 https://orcid.org/0000-0003-1882-2961 mailto:ckn@geus.dk knudsen et al. 2021: geus bulletin 47. 5336. https://doi.org/10.34194/geusb.v47.5336 2 of 11 www.geusbul let in.org saltwater intrusion may be the most important challenge in coastal aquifers, and there is a strong need for the development and implementation of effective mitigation and remediation measures (hinsby et al. 2018). designing efficient mitigation and remediation measures requires an understanding of the sources and occurrence of the saltwater. possible sources include saltwater from nearby seas, marine deposits in or adjacent to the aquifers or saline groundwater recharge due to decreasing and/or increasing precipitation and evaporation. all these sources may affect the salinity of the aquifers in some areas, while in others perhaps only one of the sources is of significance. the final decision on which method should be implemented to protect coastal water resources ultimately relies on a sound understanding of the location and nature of the salinisation sources. here, we propose a new approach based on self-organising maps (soms) for fingerprinting the sources of increasing salinity in aquifers using an example from a coastal confined chalk aquifer on southern falster, denmark (fig. 1a). elevated salinity levels were observed at shallow depth in a maastrichtian chalk aquifer on the island of falster, south-eastern denmark (fig. 1) as early as 1936 (ødum & christensen 1936). chloride concentrations of c. 200–300 mg/l cl– are locally found down to c. 25 m below the surface (m b.s.). the chloride content is generally below the who/eu drinking-water standard of 250 mg/l cl– but high compared with the recent groundwater recharge and frequently above a threshold value of c. 200 mg/l cl– proposed for danish groundwater bodies based on the european water framework and groundwater directives and related guidelines (hinsby et al. 2008). below this, between 25 and 45 m b.s., the salinity is c. 500 mg/l cl– and it increases with depth to c. 4000 mg/l cl– at 80 m depth and further to c. 20 000 mg/l cl– at about 150 m b.s. (rasmussen et al. 2013). a recent status assessment of danish groundwater bodies has identified the chalk aquifer on falster as having poor chemical and quantitative status due to indications of freshwater mining and saltwater intrusion in several areas (henriksen et al. 2019). this agrees with a previous assessment, which concluded that new innovative subsurface water solutions involving managed aquifer recharge (zuurbier et al. 2017) will most probably be required to secure future sustainable water supply at least in the southern part of falster (hinsby et al. 2018). the objective of this study was to investigate the origin of the elevated salinity in the area: is it caused by intruding saltwater from the baltic sea, up-coning of salt connate water (water trapped in the pores of the chalk when deposited) from below due to pumping, incomplete leaching of connate water in the upper part of the chalk aquifer or by infiltration of recharge through holocene marine sediments above the chalk aquifer (hinsby et al. 2012; rasmussen et al. 2013)? understanding the sources of the saltwater have implications for the choice of measures implemented to control further increases of chloride in water supply wells in the area (e.g. zuurbier et al. 2017). to understand the sources of chloride (saltwater) in water-supply wells, we analysed (1) water from the 50 km eocene–miocene paleocene above danian danian limestone upper cretaceous chalk older than upper cretaceous u denmark sweden aa cc dd ee ssssi a g 2 km gg ll oo ccii ttee cctt oonn ii pprr ee oonn mm rr iinn aall mm aa oorr test site ba iinnee i a g fig. 1 location and geological setting of the study site. a: geology of denmark. b: focussed view of the test site on the island of falster, south-eastern denmark. the location of the test site (fig. 3) is indicated by a red circle. modified from pedersen et al. (2018). https://doi.org/10.34194/geusb.v47.5336 http://www.geusbulletin.org knudsen et al. 2021: geus bulletin 47. 5336. https://doi.org/10.34194/geusb.v47.5336 3 of 11 www.geusbul let in.org baltic sea, (2) connate water from deeper levels in the chalk, (3) water from filters in wells in the chalk at depths from where the water company abstracts drinking water and (4) the variation of the composition over time while pumping on the wells. the main aquifer in the area consists of maastrichtian chalk in which the quaternary glaciations have caused intense fracturing in the upper 10–43 m of the chalk. the glacial disturbances include dislocation and development of a chalk-glaciotectonite locally containing quaternary sediments (pedersen et al. 2018). below 45 m b.s., the gamma-ray signal increases suggesting that the chalk is marly and the transmissivity in the chalk is lower here. in danish chalk aquifers, the residual saltwater is usually completely flushed out in the upper 50–80 m (bonnesen et al. 2009). below this follows a transitional zone with increasing salinity to depths of 150–200 m where the chloride concentration exceeds 19  000 mg/l, similar to deeper parts of the chalk in, for example, england (edmunds et al. 1987). this level of salinity is similar to oceanic seawater and may represent connate water. the low-lying glacial depression south-east of the marginal moraine (fig. 1b) formed a shallow lagoon in the holocene until land reclamation was initiated in the early 1860s, and it became operational from around 1900. the area is now covered by a thin layer of holocene marine sand (figs 2, 3b). 2 methods 2.1 sampling strategy groundwater is pumped from well ub2 and re-injected in well ub1 creating a dipole flow system between the two wells (fig. 4). the tracer (nacl) is injected in well t2, and the breakthrough is observed in samples collected from the 14 screens in wells cmt 1 and 2. the location of these wells is marked on fig. 4. trace elements (measured by icp-ms, details given) and chloride (measured by ion chromatography) were analysed in samples collected from cmt 1 and 2 wells. electrical conductivity was continuously monitored in the same wells by wtw instruments using the calibrated wtw electrodes. the aim of the tracer test was to assess the dual porosity characteristics of the chalk. 2.2 sample collection and analysis two closely spaced multi-level monitoring wells (cmt 1 and 2), about 20 m downstream of a water supply well (242.212), were sampled in the maastrichtian chalk aquifer (fig. 4). a total of 58 ground water samples were taken from 12 1-m screens, between 14 and 26 m b.s. the samples were collected during continuous pumping on the wells between 1 may and 4 june 2018. in the period from 1 may to 22 may 2018, 4.9 m3/h were pumped from a well (ub2) 3 m downstream (in the direction of the regional flow during the tracer test) from 0 5 10 15 20 25 30 35 d ep th b el ow s ur fa ce (m ) dgu well no 242.394 (ub3) dgu well 242.395 0 5 10 15 20 marine sand soil peat and gyttja glacio�uval sand clayey lodgement till cretaceous chalk flint in chalk chalk-glacitectonite 21 d ep th b el ow s ur fa ce (m ) b 242.394 (ub3) 242.384 (ub1) 18 31 35 43 10 d ep th b el ow s ur fa ce (m ) c on ce al ed m el tw at er c ha nn el décollement zone chalk bedrock cover of glacial deposits on chalk top surface chalk glacitectonite w e breciated chalk 20 m overburden of postglacial marine sand a fig. 2 a: block diagram of the test site. note that a significant part of the tracer tests described in this study takes place in the glacitectonite between 15 and 20 m b.s. (modified after pedersen et al. 2018). two wells are shown: dgu well no. 242.394 (ub3) and 242.384 (ub1). b: schematic geological logs in two wells, dgu well no. 242.394 (ub3) and 242.395. https://doi.org/10.34194/geusb.v47.5336 http://www.geusbulletin.org knudsen et al. 2021: geus bulletin 47. 5336. https://doi.org/10.34194/geusb.v47.5336 4 of 11 www.geusbul let in.org the monitoring wells (cmt 1 and 2), and 2.6 m3/h were re-injected into the well (ub1) c. 30 m upstream from the monitoring wells (fig. 4). the constant pumping created a constant flow field during the dipole tracer test passing through the monitoring wells, such that the water sampled from a given sampling screen should represent the same volume of the aquifer for the pumping period. from 22 may to 4 june 2018, the pumping was set to 9.1 m3/h, and the re-injection was set to 5.3 m3/h. a further five samples were collected from two other wells in the area about halfway through the experiment (17–18 may): three samples representing marine water from the baltic sea, one sample from 40 m b.s. (well site dgu 242.375, 100 m from the baltic sea; fig. 3) and one sample representing deeper groundwater from 85 m b.s. (well site dgu 242.344, 2.8 km from the baltic sea; fig. 3b). the purpose of this sampling was to obtain the composition of possible end-members for comparison. all samples for the analysis of trace elements were collected by syringes and injected directly into prepared and acidified 20 ml polyethylene vials through 0.45 µm filters and kept at c. 8° celsius until analysis. the samples from the multi-level wells (cmt 1 and 2) were collected by syringes directly from the 4 mm tubing constantly discharging water from the different screens at c. 80 ml per min. the samples were analysed for trace elements at geus using inductively coupled plasma mass spectrometry (icp-ms) and the perkin–elmer ‘totalquant’ method. methods are provided in supplementary file s1, online. we analysed for the following elements: al, as, b, ba, ca, ce, co, cr, cs, cu, fe, k, la, li, mg, mn, mo, na, ni, p, rb, sc, se, si, sr, th, ti, u, v, y and zn (the data are available online in supplementary file s2). 2.3 self-organising map analysis the unsupervised machine-learning algorithm, self-organising map (som), provides a non-linear method for visualising multi-dimensional data, and it is used here to analyse the geochemical data. the som may holocene marine sand cross section a–a’ a’a w 242.128 elevation m a.s.l. 242.239 242.375 se pleistocene sand pleistocene clay till chalk 500 1000 1500 2000 2500 3000 3500 40000 0 10 –10 –20 –30 –40 well no. 242.344 well no: 242.128 well no. 242.212 well no. 244.375 well no. 242.239 1 km map data ©2021 google marielyst væggerløse a a’ 2021 lagoon (bøtø nor) barrier islands push moraine 1780a b fig. 3 a: the study area in 1780, about 80 years before land reclamation was initiated. b: the study area in 2021 with drainage canals developed in the land reclamation project initiated in 1861. the tracer test site investigated in this study (fig. 4) is established at well no.: 242.212 (yellow triangle). the cross-section a–a’ through the study area and tracer test site is shown in the lower panel. the local water works abstracts groundwater from 11 wells in the upper 10 m of the chalk in the area. https://doi.org/10.34194/geusb.v47.5336 http://www.geusbulletin.org knudsen et al. 2021: geus bulletin 47. 5336. https://doi.org/10.34194/geusb.v47.5336 5 of 11 www.geusbul let in.org be viewed as a two-dimensional grid onto which the multi-dimensional input data are projected or mapped. here, we aim to reduce the complexity of our large dataset (27 elements analysed in 63 samples) into a set of geochemical fingerprints for each sample. these fingerprints can be represented in two dimensional maps, such that similar objects are close together, and dissimilar objects are far away from each other. in other words, the som processing finds an approximation of the data by mapping the input data into another dataset, which has fewer datapoints than the initial dataset. this approximation, referred to as the best matching units (bmus), has the same amount of data types (i.e. same dimensionality) as the input data. the bmus are presented in the som as a two-dimensional map and referred to as the som-space. each bmu is associated with a cell in the matrix or somspace, and they are ordered in the matrix such that similar bmus are adjacent to each other in the two-dimensional map. although the som is two-dimensional, the multi-dimensionality of the input data is retained by the dimensionality of the bmu. each of the input datapoints has a bmu to which the data are most similar. several input data may be associated with the same bmu, in which case the som presentation may be viewed as a classification of the input data. the som is discretised in a pre-selected number of rows and columns (grid or matrix representation). thus, the original multi-dimensional data are mapped into a new data space with only two dimensions and fewer data. in our case, the number of rows and columns are 5 and 11, respectively (fig. 5a). the number of cells or elements defines the data reduction applied to the initial data. in this case, a reduction from 63 data samples analysed for 27 elements = 1701 datapoints reduced to a maximum of 5 × 11 = 55 datapoints (fig. 5). the grid size is determined based on a heuristic approach. large maps produce a large number of small but ‘compact’ clusters (records assigned to each cluster are quite similar). small maps produce fewer but more generalised clusters. a ‘right number of clusters’ does not exist, especially in real-world datasets. it all depends on the detail in which you examine your dataset. the grid size used in this study is calculated by som based on the size of the dataset. the som analysis is based on the commercial software package sirosom™. the grid size used in this article is calculated by sirosom based on the size of the dataset. the bmu colours are generated by sirosom, and the timeline and cross-plots are generated using a matlab script developed by the geological survey of denmark and greenland (geus). the procedure used here is further described in supplementary file s1. the positions of the datapoints (fig. 5b) represent the analytical results, whereas the colours of the dots represent their position in the entire dataset. the colour scale is chosen by the algorithm, but the colours are chosen so that points with colours close to each other represent ub1 t1 242.212 t2 ub3 ub2cmt 1 2 0 m 10 m 20 m 30 m 40 m sand chalk till clay peat chalk with increasing saltwater water table 5 m fig. 4 design of the tracer test site and wells (ub1, t1, t2, ub3, cmt1 and 2 and ub2). red colours from about 30 m b.s. indicate high salinity with depth. the supply well is well 242.212, location on fig 3. green arrow indicates direction of the flow during the tracer test. https://doi.org/10.34194/geusb.v47.5336 http://www.geusbulletin.org knudsen et al. 2021: geus bulletin 47. 5336. https://doi.org/10.34194/geusb.v47.5336 6 of 11 www.geusbul let in.org samples with compositions that are close to each other in the 27-dimensional space. detailed information on som can be found in kohonen (1982, 1990, 2001, 2013) and kalteh et al. (2008). the som method has recently been used in the analysis of groundwater geochemical data in a coastal aquifer near venice (dalla libera et al. 2020) and in geochemical pattern recognitions of deep thermal groundwater in south korea (kim et al. 2020). 3 results the arrival of the tracer is seen as an increase in the conductivity in the monitoring wells (fig. 6). the timing varies among the three different levels with maximum effect c. 100 h after the start of the experiment at 15.5 m, whereas the tracer peak arrives later at the 19.5 and 23.5 m levels and is much more diffuse. another injection of tracer was initiated on 23 may 2018 peaking on 25 may at 15.5 m. the increased pumping and reinjection rate effects the time it takes for the tracer to pass the monitoring well (fig. 6). the gap in the measurements of electrical conductivity (ec) in fig. 6 was caused by power loss on the ec meter. 3.1 hydro-chemical fingerprints vary both with stratigraphy and time concentrations of, for example, na, mg, mn, co, sr and zn vary with depth in the wells and with time during the conducted tracer test, 1 may to 4 june 2018 (fig. 7). for the first c. 75 h, the composition of the water is rather constant at the three levels (fig. 7). the contents of na, mg, sr, zn, mn and co are high at the 19.5 m level relative to the 15.5 and 23.5 m levels and distinctly different from both the 15.5 and the 23.5 levels. the contents of mg, sr and zn at the 15.5 m level are low, and the content of mn is high relative to the 23.5 m level; the compositions of the water at the different stratigraphic levels are all distinct from each other. the content of na varies among the different stratigraphic levels (fig. 7). the content of mg is relatively low in the upper (15.5 m) level compared with both the middle (19.5 m) and lower (23.5 m) level, and the concentration stays more or less constant over the duration of the experiment, presumably because it is part of the carbonate rock itself. the content of mn is relatively high in the upper and middle level compared with the lower level with a slight decrease over time. with respect to co, the middle level is high relative to both the upper and lower levels and also shows a slight decrease over time. the sr content is high in the middle level, low in the upper level and intermediate in the lower level, and the sr content seems to be fairly stable in all levels. the content of zn is low in the upper level relative to both the middle and lower levels. collectively, the contents of na, mg, mn, co, sr and zn make the composition of each level distinct for the whole duration of the experiment. the composition of the water when taking the complete analysis (all 27 elements) into account is reflected by the colour of the dots in fig. 7. in the upper level (15.5 m), the dots range from purple and blue colours to blue–green during the first c. 80 h. then, the composition is fairly stable (blue-green colours) for the next 300 h after which it changes to compositions dominated by green colours. in the middle level (19.5 m), the compositions are marked with pink and purple colours in the first c. 80 h; after this, the compositions are marked blue for the next 300 h before turning green as is seen in the upper (15.5 m) level. at the lower level (23.5 m), the compositions are dominantly marked fig. 5 self-organising maps (soms): a: each original datapoint is associated with a best matching units (bmu) point and only 26 out of 55 bmu points are selected. no units. b: example som for mg (ppm) versus time: the original datapoints with the associated bmu colours and numbers are displayed. 21 2 1 3 4 5 6 7 8 9 10 11 22 3 4 5 11 10 32 43 54 33 55 20 42 539 19 30 187 17 39 50 16 49 3 16 7 7 7 7 30 18 19 30 9 17 19 5 3 1 15 26 1000 24 26 28 30 32 44 200 300 400 500 600 700 hours from start pp m m g ba https://doi.org/10.34194/geusb.v47.5336 http://www.geusbulletin.org knudsen et al. 2021: geus bulletin 47. 5336. https://doi.org/10.34194/geusb.v47.5336 7 of 11 www.geusbul let in.org with light brown and brown colours for the first c. 80 h. after which green colours also dominate. 3.2 comparison with sea water and water from deeper levels the composition of the water from the monitoring wells is compared with the composition of water from the baltic sea as and water from two deeper wells in the area (fig. 8). the mg concentrations of the seawater and the deeper wells are generally high compared with the monitoring wells (fig. 8). furthermore, the proportions (ratios) between the elements are different from the monitoring wells. an example is the very high mn/mg, co/mg, zn/mg and sr/mg ratios in the monitoring wells compared with the seawater (yellow dots in fig. 8). there is a linear relationship between sr and mg in the monitoring wells (grey on fig. 8). the deep well with a sample from 85 m b.s. (dgu 242.344) has a high chloride content and most trace-element contents are also high (light green dot in fig. 6). the sample from 40 m b.s. also has elevated contents of trace elements (yellow dot in grey circle in fig. 8), and the sr/mg, co/mg and zn/mg ratios are more like those in seawater samples. however, the mn/mg ratio is different. the hypothetical mixing lines between the infiltrating water and the deep, saline connate water (light green) and baltic sea water (yellow) are shown in fig. 8. the composition of the infiltrating water is inferred to be rainwater containing c. 1% baltic sea water due to the near-coastal position and influx of wind-carried salty aerosols. this corresponds to a chloride concentration of about 40 mg/l, which has been observed in three nearby water-supply wells in chalk about 1 km north-west of the tracer test site, which seem to have been completely freshened during the holocene. the composition of the infiltrating water is indicated by the trend lines on fig. 8. the water sample from 40 m (marked with a grey ring) has the same colour as the dots representing the seawater (yellow; fig. 8). this implies that their hydro-chemical fingerprints are alike, that is, they are very close to each other in the 27-dimensional space of the geochemical dataset. 4 discussion trace-element analysis of freshwater has previously been tested as a method for understanding the water– rock interactions in a system. schürch et al. (2004) suggested that the groundwater geochemistry in a chalk aquifer is dominated by incongruent reactions with the fine-grained carbonate sediments, which release trace-element impurities into the water. khadra et al. (2017) found that the chemical differences between limestone and dolomitic limestone were important factors in explaining the observed variation in the concentration of trace elements in fresh groundwater. fig. 6 electrical conductivity (µs/cm) of the water in the monitoring wells as a function of time. the nacl tracer was injected twice during the period on 1 may 2018 at 16.00 and 23 may 2018 at 14.24, indicated by green and purple vertical lines. abstraction from ub2 was increased from 4.9 to 9.1 m3/h during the second tracer test resulting in a faster breakthrough. 1000 1100 1200 1300 1400 1500 1600 –24 72 168 264 360 456 552 648 744 840 el ec tr ic al c on du ct iv it y µs /c m 0 time since 1st injection – hours 15.5 mbs 19.5 mbs 23.5 mbs 1st tracer injection 2nd tracer injection https://doi.org/10.34194/geusb.v47.5336 http://www.geusbulletin.org knudsen et al. 2021: geus bulletin 47. 5336. https://doi.org/10.34194/geusb.v47.5336 8 of 11 www.geusbul let in.org 100 10-3 10-2 0 200 300 hours from start 400 500 600 1000 200 300 400 500 600 1000 200 300 400 500 600 hours from start hours from start 15.5 m 19.5 m 23.5 m zn 2.6 3.2 3.4 2.8 3 2.4 sr 3 6 8 7 4 5 2 c o 6 9 11 12 10 7 8 5 28 36 38 34 30 32 26 24 m n m g 90 130 120 100 110 80 70 n a fig. 7 concentrations (ppm) of na, mg, mn, co, sr and zn in the extracted water from the three filters at 15.5 m, 19.5 m and 23.5 m as a function of time (hours) from 1 may to 4 june 2018. depth of sampling points is shown at the top. note that the y-axis is on a log scale apart from mg. the position of the points represents the actual measurements, whereas the colour of the points represents the som derived geochemical fingerprint. https://doi.org/10.34194/geusb.v47.5336 http://www.geusbulletin.org knudsen et al. 2021: geus bulletin 47. 5336. https://doi.org/10.34194/geusb.v47.5336 9 of 11 www.geusbul let in.org yuan et al. (2019) found that multi-element fingerprinting can be useful for assessing hydrological connectivity across the landscape and indicate that element concentrations are affected not only by land use but also by hydrogeological characteristics. the trace-element composition of the groundwater at 15.5, 19.5 and 23.5 m b.s. (fig. 7) is distinct from each other at early times in the experiment, before the composition of the water is homogenised by the recirculation. apart from na, these elements occur in different minerals in the chalk with mg and sr mainly in carbonates, zn and co mainly in sulphides and mn mainly in oxides. the content, composition and/or reactivity of these minerals presumably vary among the different stratigraphic levels. the most likely explanation for the variation and distinct chemistry of the water at the different stratigraphic levels at the start of the experiment is that the different trace-element distributions in the water reflect the variable composition of the chalk with depth, that is, the content of clay minerals, carbonates and sulphides and variable trace-element composition of these minerals. this implies that the water composition has evolved due to geochemical reactions with the chalk at various stratigraphic levels, as suggested by schürch et al. (2004). this could be due to variations in trace-element composition at the time of formation of the chalk, a result of post-depositional solution or precipitation or some other type of reaction, such as sorption of the trace elements on the mineral surfaces or variations in the contents of minerals due to glaciotectonic redistribution of sediments. however, the mechanism controlling the variation in trace-element concentrations is not clear. up-coning and seawater intrusion were suggested to be responsible for increasing salinity in wells close to each other in a similar coastal carbonate aquifer about 100 km north of the falster test site (thorn 2011). if the trace element concentrations in the aquifer were a result of simple mixing between meteoric water and baltic sea water, the samples should fall on a mixing line between the two sources. furthermore, the colour of the dots representing the compositional variation should reflect fig. 8 concentrations (ppm) of co, mn, sr and zn and versus ppm mg in samples from the pumping experiment compared with baltic seawater and two deeper wells on falster at 85 m b.s. and 40 m b.s. note that it is a log–log plot. 5 10 20 30 50 100 ppm mg 200 pp m s r 0,01 50 100 ppm mg 200 0,001 pp m z n 4 5 6 9 50 100 200 8 7 3 pp m c o * 10 4 7 8 9 12 50 100 200 11 10 5 6 pp m m n * 10 3 monitoring wells trend mixing seawater mixing deep salt water salt water (85 m) salt water (40 m) seawater https://doi.org/10.34194/geusb.v47.5336 http://www.geusbulletin.org knudsen et al. 2021: geus bulletin 47. 5336. https://doi.org/10.34194/geusb.v47.5336 10 of 11 www.geusbul let in.org this (fig. 8). however, this is not the case, and thus, the elevated salinity in the monitoring wells is unlikely to be the result of simple dilution of baltic sea water. similarly, if the trace element concentrations in the aquifer were a result of simple mixing between meteoric water and connate water from deeper levels, the samples should likewise fall on a mixing line between these two sources. this is not the case, as the trace-element composition is also distinct from the connate saline water from the deeper levels of the chalk aquifer (fig. 8), suggesting that elevated salinity in the wells is not simply caused by up-coning of saline water from below. after c. 80 h, the composition of the water in the monitoring wells changes slightly at all three levels as seen in the change of the colours of the dots marking the geochemical fingerprint of the water derived from the som analysis (fig. 7). the geochemical fingerprint of the water is very distinct at the 19.5 m level with bright pink dots and light brown dots at the 23.5 m level. at the 19.5 m level, the composition changes after c. 80 h and is marked by blue dots for the following 300 h. the composition of the water at this level is still rather different from the other two levels, and the reason for the conservative nature of the composition here may be caused by the fairly high levels of trace elements in the water. after c. 500 h, the composition at the 19.5 m level changes, reflected by green dots (fig. 7). at the 23.5 m level, the change to green dots occurs after c. 80 h, and at the 15.5 m level, the green dots occur after c. 500 h. however, the colour of the dots at the 15.5 m level becomes blue–green after c. 80 h indicating that the composition is developing towards the green dots after 500 h at 19.5 m and after 80 h at 23.5 m. at 15.5 m, the tracer concentration peaks earlier than at deeper levels (fig. 6) and seems to be flushed out earlier than the other levels. this is probably because the chalk is heavily fractured here (fig. 2b). the shift towards greenish colours over time at all levels indicates that the samples become statistically more homogeneous, which is most likely an effect of the recirculation mixing the water from all levels and smoothing out their characteristic fingerprints. this indicates that the fingerprint characteristics must be a feature that develops over time through water–rock interactions. the water analysed from 40 m b.s. in the dgu 242.375 well located 100 m from the baltic sea is very similar to the water from the baltic sea (fig. 8) with respect to the content of mg, sr and co. furthermore, the dot that reflects the overall geochemical fingerprint, taking all the 27 elements into account, has the same colour as that for seawater (yellow). this implies that the geochemical fingerprint in this well is similar to the seawater and that the water here is derived from the baltic sea. the different mn content may be caused by trapping of mn on oxides towards the well and/or the well screen. the present position of the saltwater–freshwater interface at c. 30–40 m b.s. (c. –30 to –40 m a.s.l.) is controlled by the depth and distribution of hydraulic active fractures in the chalk, as was found for sites on the island of sjælland, c. 100 km to the north of falster where the interface is found between –45 and –70 m a.s.l. (larsen et al. 2006). the thinner freshwater zone on the island of falster might be explained by a number of processes, including the ablation history of retreating glaciers during past glaciations, the marine setting of the aquifer during most of the holocene and finally more recent land reclamation on falster, which resulted in very low groundwater recharge since the early 1860s. multi-element fingerprinting of water is not commonly used in northern europe, and the results of the present investigation suggest that this can be a useful method to investigate water–rock interactions, water provenance and the sources of dissolved chloride. this information is important when planning measures to control saltwater intrusion and protect drinking water resources such as managed aquifer recharge of desalinised brackish groundwater from deeper parts of the aquifer. the results of this study and of previous studies by rasmussen et al. (2013), pedersen et al. (2018) and hinsby et al. (2018) clearly demonstrate that a sound understanding of local geological and geochemical features of aquifers is a prerequisite for the protection of water-supply wells from saltwater intrusion of various sources in coastal areas. 5 conclusions water extracted from a slightly saline aquifer on falster has distinctly different hydro-chemical fingerprints depending on the stratigraphic levels from where it is extracted. this suggests that the composition has evolved by interaction with the rock. the slightly saline character of the water is probably due to incomplete refreshing of the aquifer during the holocene. this could again be caused by the very low relief and flow in the area. water composition changes with time during the tracer experiment, probably caused by the flushing of the reservoir with re-injected water. introduction of the (nacl) tracer may also have caused ion exchange, thus mobilising other cations in the reservoir. comparisons of aquifer water with baltic sea water and deep connate water show that the observed compositions cannot be generated by mixing these sources with the meteoric recharge. however, the hydro-chemical composition/fingerprint of the water in another well extracted at 40 m b.s. is very close to seawater, https://doi.org/10.34194/geusb.v47.5336 http://www.geusbulletin.org knudsen et al. 2021: geus bulletin 47. 5336. https://doi.org/10.34194/geusb.v47.5336 11 of 11 www.geusbul let in.org suggesting that the water here contains a high fraction of seawater. this is likely because the well is located close to (c. 100 m from) the baltic sea. the statistical analysis, based on an unsupervised machine-learning algorithm, som, has proved to be a very effective way to facilitate interpretation of a large number of variables and thus understand the compositional variation in an aquifer. using multi-element analysis combined with som analysis/fingerprinting may also be useful in future studies of variations in groundwater quality and, for example, in characterising the source of chloride in coastal aquifers. acknowledgements the authors would like to thank olga nielsen (geus) for the analytical work, per jensen (geus) for help during the fieldwork and peer jørgensen (copenhagen university) for the development of the partly automated sampling and monitoring system in cmt wells. additional information funding the study was conducted as part of the subsol project (www.subsol. org) that received funding from the european union’s horizon 2020 research and innovation programme under grant agreement no. 642228. author contributions all: investigation and writing (original draft, review and editing). ck and ljk: data curation and methodology. kh: project administration, conceptualisation and funding acquisition. rj: conceptualisation and methodology. pr: conceptualisation. additional files two additional files are available at: https://doi.org/10.22008/fk2/ q8yf8v references bonnesen, e.p., larsen, f., sonnenborg, t.o., klitten, k. & stemmerik, l. 2009: deep saltwater in chalk of north-west europe: origin, interface characteristics and development over geological time, hydrogeology journal 17, 1643–1663. https://doi.org/10.1007/s10040-009-0456-9 dalla libera, n., pedretti, d., tateo, f., mason, l., piccinini, l. & fabbri, p. 2020: conceptual model of arsenic mobility in the shallow alluvial aquifers near venice (italy) elucidated through machine learning and geochemical modeling. water resources research 56, e2019wr026234. https://doi.org/10.1029/2019wr026234 edmunds, w.m., cook, j.m., darling, w.g., kinniburgh, d.g., miles, d.l., bath, a.h., morgan-jones, m. & andrews, j.n. 1987: baseline geochemical conditions in the chalk aquifer, berkshire, u.k.: a basis for groundwater quality management. applied geochemistry 2, 251–274. https://doi.org/10.1016/0883-2927(87)90042-4 henriksen, h.j., voutchkova, d., troldborg, l., ondracek, m., schullehner, j. & hansen, b. 2019: national vandressource model beregning af udnyttelsesgrader, afsænkning og vandløbspåvirkning med dk model 2019. geus report 2019/32, 84 pp. hinsby, k., condesso de melo, m.t. & dahl, m. 2008: european case studies supporting the derivation of natural background levels and groundwater threshold values for the protection of dependent ecosystems and human health. science of the total environment 401, 1–20. https://doi.org/10.1016/j.scitotenv.2008.03.018 hinsby, k., jessen, s., larsen, f. & postma, d. 2012: a groundwater chemistry and multi-tracer study of sources of salt water intrusion – the island of falster, denmark. 22nd salt water intrusion meeting, swim2012, 18–22 june, buzios, brazil. hinsby, k., johnsen, a.r., rasmussen, p., sonnenborg, t.o., sørensen, s.r., postma, d., thorn, p., scharling, p.b. & gudbjerg, j. 2016: water4coasts – new methods for integrated management and protection of coastal aquifers, 160 pp. copenhagen: danish ministry of environment and food. hinsby, k. et al. 2018: d2.2 road map for full scale implementation of sws for fractured chalk aquifer. 68 pp. deliverable for the eu project subsol. www.subsol.org kalteh, a.m., hjorth, p. & berndtsson, r. 2008: review of the self-organizing map (som) approach in water resources: analysis, modelling and application. environmental modelling & software 23, 835–845. https:// doi.org/10.1016/j.envsoft.2007.10.001 khadra, w.m., stuyfzand, p.j. & van breukelen, b.m. 2017: hydrochemical effects of saltwater intrusion in a limestone and dolomitic limestone aquifer in lebanon. applied geochemistry 79, 36–51. https:// doi.org/10.1016/j.apgeochem.2017.02.005 kim, k.h., yun, s.t., yu, s., choi, b.y., kim, m.j. & lee, k.j. 2020: geochemical pattern recognitions of deep thermal groundwater in south korea using self-organizing map: identified pathways of geochemical reaction and mixing. journal of hydrology 589, 125202. https://doi. org/10.1016/j.jhydrol.2020.125202 kohonen, t. 1982: self-organized formation of topologically correct feature 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important lithology in former glaciated terrains covering chalk and limestone bedrock. geological survey of denmark and greenland bulletin 41, 21–24. https://doi.org/10.34194/geusb.v41.4333 rasmussen, p., sonnenborg, t.o., goncear, g. & hinsby, k. 2013: assessing impacts of climate change, sea level rise, and drainage canals on saltwater intrusion to coastal aquifer. hydrology and earth system sciences 17, 421–443. https://doi.org/10.5194/hess-17-421-2013 schürch, m., edmunds, w.m. & buckley, d. 2004: three-dimensional flow and trace metal mobility in shallow chalk groundwater, dorset, united kingdom. journal of hydrology 292, 229–248. https://doi. org/10.1016/j.jhydrol.2004.01.004 thorn, p. 2011: groundwater salinity in greve, denmark: determining the source from historical data. hydrogeology journal 19(2), 445–461. https://doi.org/10.1007/s10040-010-0680-3 yuan, y., zhu, x., mushet, d.m. & otte, m.l. 2019: multi-element fingerprinting of waters to evaluate connectivity among depressional wetlands. ecological indicators 97, 398–409. https://doi.org/10.1016/j. ecolind.2018.10.033 zuurbier, k., raat, k.j., paalman, m., oosterhof, a.t. & stuyfzand, p.j. 2017: how subsurface water technologies (swt) can provide robust, effective, and cost-efficient solutions for freshwater management in coastal zones. water resources management 31, 671–687. https://doi. org/10.1007/s11269-016-1294-x https://doi.org/10.34194/geusb.v47.5336 http://www.geusbulletin.org http://www.subsol.org http://www.subsol.org https://doi.org/10.22008/fk2/q8yf8v https://doi.org/10.22008/fk2/q8yf8v https://doi.org/10.1007/s10040-009-0456-9 https://doi.org/10.1029/2019wr026234 https://doi.org/10.1016/0883-2927(87)90042-4 https://doi.org/10.1016/j.scitotenv.2008.03.018 http://www.subsol.org https://doi.org/10.1016/j.envsoft.2007.10.001 https://doi.org/10.1016/j.envsoft.2007.10.001 https://doi.org/10.1016/j.apgeochem.2017.02.005 https://doi.org/10.1016/j.apgeochem.2017.02.005 https://doi.org/10.1016/j.jhydrol.2020.125202 https://doi.org/10.1016/j.jhydrol.2020.125202 https://doi.org/10.1007/bf00337288 https://doi.org/10.1007/bf00337288 https://doi.org/10.1109/5.58325 https://doi.org/10.1007/978-3-642-56927-2 https://doi.org/10.1016/j.neunet.2012.09.018 https://doi.org/10.34194/geusb.v41.4333 https://doi.org/10.5194/hess-17-421-2013 https://doi.org/10.1016/j.jhydrol.2004.01.004 https://doi.org/10.1016/j.jhydrol.2004.01.004 https://doi.org/10.1007/s10040-010-0680-3 https://doi.org/10.1016/j.ecolind.2018.10.033 https://doi.org/10.1016/j.ecolind.2018.10.033 https://doi.org/10.1007/s11269-016-1294-x https://doi.org/10.1007/s11269-016-1294-x fingerprinting sources of salinity in a coastal chalk aquifer in denmark using trace elements abstract 1 introduction 2 methods 2.1 sampling strategy 2.2 sample collection and analysis 2.3 self-organising map analysis 3 results 3.1 hydro-chemical fingerprints vary both with stratigraphy and time 3.2 comparison with sea water and water from deeper levels 4 discussion 5 conclusions acknowledgements additional information references figures fig. 1 location and geological setting of the study site. a: geology of denmark. b: focussed view of the test site on the island of falster, south-eastern denmark. the location of the test site (fig. 3) is indicated by a red circle. modified from pedersen et al. (2018). fig. 2 a: block diagram of the test site. note that a significant part of the tracer tests described in this study takes place in the glacitectonite between 15 and 20 m b.s. (modified after pedersen et al. 2018). two wells are shown: dgu well no. 242.394 (ub3) and 242.384 (ub1). b: schematic geological logs in two wells, dgu well no. 242.394 (ub3) and 242.395. fig. 3 a: the study area in 1780, about 80 years before land reclamation was initiated. b: the study area in 2021 with drainage canals developed in the land reclamation project initiated in 1861. the tracer test site investigated in this study (fig. 4) is established at well no.: 242.212 (yellow triangle). the cross-section a–a’ through the study area and tracer test site is shown in the lower panel. the local water works abstracts groundwater from 11 wells in the upper 10 m of the chalk in the area. fig. 4 design of the tracer test site and wells (ub1, t1, t2, ub3, cmt1 and 2 and ub2). red colours from about 30 m b.s. indicate high salinity with depth. the supply well is well 242.212, location on fig 3. green arrow indicates direction of the flow during the tracer test. fig. 5 self-organising maps (soms): a: each original datapoint is associated with a best matching units (bmu) point and only 26 out of 55 bmu points are selected. no units. b: example som for mg (ppm) versus time: the original datapoints with the associated bmu colours and numbers are displayed. fig. 6 electrical conductivity (μs/cm) of the water in the monitoring wells as a function of time. the nacl tracer was injected twice during the period on 1 may 2018 at 16.00 and 23 may 2018 at 14.24, indicated by green and purple vertical lines. abstraction from ub2 was increased from 4.9 to 9.1 m3/h during the second tracer test resulting in a faster breakthrough. fig. 7 concentrations (ppm) of na, mg, mn, co, sr and zn in the extracted water from the three filters at 15.5 m, 19.5 m and 23.5 m as a function of time (hours) from 1 may to 4 june 2018. depth of sampling points is shown at the top. note that the y-axis is on a log scale apart from mg. the position of the points represents the actual measurements, whereas the colour of the points represents the som derived geochemical fingerprint. fig. 8 concentrations (ppm) of co, mn, sr and zn and versus ppm mg in samples from the pumping experiment compared with baltic seawater and two deeper wells on falster at 85 m b.s. and 40 m b.s. note that it is a log–log plot. geological survey of denmark and greenland bulletin 12, 24-46 24 0 10 20 30 40 50 60 70 80 90 100 cm 2239 m 2240 m 2245 m 2246 m cecilie-1 fig. 10. core photographs showing dark grey, largely structureless, discordant, intrusive sandstones within the lighter grey mudstones of the upper tyr member in the cecilie-1 well. depths are core depths. stratigraphic positions of the figured intervals are shown on fig. 8. lithostratigraphy rogaland group the rogaland group was established by deegan & scull (1977) and comprises the paleocene to lower eocene marlstone and mudstone succession between the top of the ekofisk formation of the chalk group (deegan & scull 1977) and the glaucony-rich mudstones of the hordaland group (now stronsay and westray groups) in the central north sea. in most of the danish sector, the rogaland group has a relatively uniform thickness and comprises the våle, lista, sele and balder formations (fig. 7). the fur formation is a part of the rogaland group and is present in a limited area in the north-eastern part of the danish sector of the north sea stretching into the norwegian sector. hardt et al. (1989) added three new sandstone units to the rogaland group in the southern viking graben (norwegian sector): ty formation, heimdal formation and hermod formation (fig. 4). although these sandstone units are broadly comparable to coeval sandstone units encountered in the siri canyon system (figs 1, 4) in the danish sector, the norwegian units and the siri canyon sandstones have different provenances and are not contiguous with each other. therefore, the sandstone units in the danish sector are described herein as new members. in the siri canyon (fig. 1), the mudstones of the rogaland group contain concordant or discordant postdepositional sandstone intrusions (hamberg et al. 2005). in some wells (e.g. cecilie-1 and nini-3) sandstone intrusions are very common (figs 8–11). the sands have intruded most levels in the rogaland group, but the ve member (new member of the lista formation, see below) in the middle part of the group is particularly rich in intrusions. most of the intrusive sandstones are only a few millimetres thick, but they may reach a thickness of 5 m. in some wells they constitute up to 30% of the total sandstone thickness. the intrusions are usually massive, but faint lamination is locally present, especially at the top of the beds. most of the intrusive sandstone bodies are separated by in situ mudstones, but they may also occur in intervals showing multiple intrusions. the boundaries with the host rock are slightly to very irregular or wavy, and in places discordant. minor intrusive offshoots (apophyses) into the host rock are common. the petrography of the intrusive sandstones is similar to that of the in situ sandstones and they are therefore most likely sourced from the 25 latter. the intrusion of sand was mainly subhorizontal, parallel to the bedding, and most of the intrusions can thus be classified as sills. the sandstone intrusions are either unconsolidated or cemented by calcite or other carbonate minerals. in some wells the intrusive sandstones are chlorite-cemented. in the nini-3 cores, the vile and ve members of the lista formation (new members, see below) are particularly rich in intrusive sandstones (figs 11, 12); the intrusions increase in number and thickness upward through the vile member to terminate in a large intrusion complex in the ve member. commonly, mudstone clasts are abundant in the sandstones and in the intrusion complex in the nini-3 well. where present, they constitute from a few percent up to 90% of the volume of the host sandstones. they are most abundant in the upper parts of the beds. the clasts range in size from a few millimetres to wider than the core diameter. most of the clasts are angular, often with delicate protrusions, and aligned parallel to the bounding planes of the intrusive sandstone body. in thick intrusions, flow banding and dewatering structures are occasionally present. top-bed rip-down mudstone clasts (stow & johansson 2000) are common (fig. 12). fossil wood fragments are present, but rare. as the distribution of intrusive sandstones is the result of postdepositional rather than synsedimentary processes, they may cross lithostratigraphic boundaries. when an injected sandstone body occurs in direct contact with an in situ sandstone unit (e.g. as seen in the higher parts of the new tyr member in fig. 8), it is impossible to distinguish between the two genetically different units on the basis of petrophysical logs and cuttings samples alone; only a sedimentological study of core material may reveal the different nature of the two sandstones. våle formation history. the våle formation was established by hardt et al. (1989) for the marls with interbedded claystones, limestones and siltand sandstone stringers that overlie the chalk group in the central and northern north sea. the lista fm vile mb tyr mb idun mb nini-3 clay si. vf. f. m. sand c. vc. p c c log depth core depthgr sonic 1760 1750 1740 1730 1760 1750 1740 1730 fig. 11. core log showing intrusive sandstones in the vile and idun members in the nini-3 well. for legend, see fig. 9. the two intervals marked by grey bars in the core depth column are shown as core photographs in fig. 12. 26 fig. 12. core photographs of the vile and idun members in the nini3 well showing sandstone intrusions, weak flow banding (1), injection breccia with abundant irregular and angular clasts (2), mudstone clasts in injected sand (3) and top bed rip-down clasts (4). depths are core depths. stratigraphic positions of the figured intervals are shown on fig. 11. fig. 13. e-8, danish reference well for the våle and lista formations, and reference well for the vile, ve and bue members. black bar shows cored section. 1726 m 1727 m nini-3 1739 m 0 10 20 30 40 50 60 70 80 90 100 cm 4 3 1 2 2060.3 2044.0 2030.3 2027.6 2057.0 horda fm balder fm sele fm lista fm bue mb ve mb vile mb våle fm chalk gp 2100 2000 m e-8 gr sonic presence of a marly succession on top of the chalk group was previously noted by deegan & scull (1977) and treated informally as an equivalent to the more coarse-grained maureen formation in the uk sector of the north sea. kristoffersen & bang (1982) established the north sea marl for an exclusively marly and calcareous unit corresponding to the maureen formation-equivalent unit of deegan & scull (1977). although the description of the north sea marl fulfils the requirements for a formal description of a lithostratigraphic unit (with the exception of lacking indication of the rank of the unit), they specifically stated that their unit was only informally established. the name north sea marl has only rarely been used outside the danish sector of the north sea; the sediments are instead referred to the våle formation, which covers most national sectors of the north sea basin. as doubt may be raised about the formal status of the north sea marl unit, and in order to promote communication between north sea stratigraphers, it is considered by the present authors that the våle formation of hardt et al. (1989) serves as the better name for the marlstone unit. type well. norwegian sector well 1/3-1, 3258–3209 m mdkb. danish reference wells. e-8, 2060.3–2057.0 m mdkb (fig. 13);siri-1,2186.5–2156.3mmdkb(fig.14;plates1, 4). distribution and thickness. the våle formation and its equivalents are present throughout the north sea basin, except in a few areas where their absence is due to nondeposition or erosion. the våle formation is absent on 27 pyrite-bearing marlstones dominate the formation (fig. 16). thin sandstone intrusions are present locally. in the siri canyon, the marls are interbedded with turbidite sandstones; where sandstone-dominated, the succession is referred to a new member (bor member, defined below). log characteristics. from its base to its top, the våle formation is characterised by an overall steady increase in gamma-ray response, combined with an overall steady decrease in sonic readings. when the bor member sandstones are present, blocky log signatures with higher gamma-ray values and lower sonic readings interrupt this general trend (fig. 17). boundaries. in most wells in the danish sector, the change from the chalks of the chalk group to the marlstones of the våle formation is gradational and the boundary can be difficult to position (fig. 16). in the siri canyon, however, most wells show an erosional contact between the chalk group and the våle formation and the formation boundary is sharp. on the petrophysical logs, the boundary is placed where the stable, low gamma-ray response characteristic of the ekofisk formation starts to increase upwards and the high sonic readings (also characteristic of the latter formation) start to decrease upwards. the change in the log pattern may be stepwise with each step represented by a small increase in gamma-ray values and an accompanying decrease in sonic readings. the våle formation is overlain by the lista formation. intrabasinal highs (hardt et al. 1989) and in parts of the siri canyon where the rogaland group overlies the chalk group with an erosionalunconformity. its thickness varies from 0 to 48 m in the danish sector of the north sea (fig. 15). lithology. light grey to greenish grey, heavily bioturbated fig. 15. isochore map of the våle formation in the study area. the positions of the two danish reference wells, e-8 and siri-1, are indicated on the figure. 2047.5 2072.6 2186.5 2156.3 horda fm balder fm sele fm lista fm rind mb idun mb tyr mb vile mb ve mb bue mb bue mb våle fm chalk gp 2000 2100 2200 m siri-1 gr sonic neutron/densityfig. 14. siri-1, danish reference well for the våle and sele formations. black bars show cored sections. e-8 siri-1 0 10 20 30 40 thickness (m) våle formation 25 km 28 cm 2398 m 2399 m 2400 m cecilie-1b 2401 m 2402 m 0 10 20 30 40 50 60 70 80 90 100 cm underlying ekofisk formation is characterised by the ho of the dinoflagellate senoniasphaera inornata followed uphole by the ho of the planktonic foraminifer globoconusa daubjergensis. there is a hiatus at the contact between the ekofisk and våle formation in many sections and wells (clemmensen & thomsen 2005). the basal part of the våle formation, just above the top of the ekofisk formation, is marked by the downhole increase in calcareous foraminifer diversity and the ho of the planktonic foraminifers globanomalina cf. compressa and subbotina subdivision. the våle formation includes a sandstone unit (bor member, new) in the danish north sea sector. macroand ichnofossils. fragments of shelly macrofossils are present, but rare. the våle formation is heavily bioturbated. trace fossils in the formation include chondrites ispp., phycosiphon ispp., planolites ispp. and zoophycos ispp. thalassinoides ispp. burrows are only present locally. microfossils and palynomorphs. the uppermost part of the fig. 16. core photographs of the ekofisk–våle formation boundary in the cecilie-1b well. the shift from chalk to marlstones is gradational and placing the boundary can be difficult; it is positioned in the middle part of the core interval 2400.00–2401 m, at 2400.35 m (arrow), where light grey marls become dominant. depths are core depths. 29 trivialis. thus, the boundary between the two formations may in practice be located by reference to these three events (fig. 5a). the ho of the dinoflagellate alisocysta reticulata marks a level in the lower part of the våle formation. calcareous microfossil events near the top of the våle formation in the danish sector include the provisional ho of planktonic foraminifers. depositional environment. over most of the danish sector, the marlstones of the våle formation comprise hemipelagic deposits and deposits from dilute turbidity currents. the marlstones are probably largely of turbiditic origin. the foraminifer fauna of the våle formation is characterised by common calcareous taxa. taxa belonging to the neritic ‘midway-type’ fauna (berggren & aubert 1975) are especially common. the plankton/benthos ratio varies from approximately 1:1 in some areas to a total dominance of calcareous benthic foraminifers in other areas. the microfaunal composition indicates that the våle formation was deposited in an open marine, outer neritic environment that periodically reached upper bathyal depths. the bottom conditions were predominantly oxic with periods of dysoxia. the indications from the microfauna are supported by the trace fossil assemblage, which indicates water depths of at least 200 m combined with oxic to dysoxic bottom conditions. in the siri canyon, where thin turbidites are common in the våle formation, gravity flows played a major role during the deposition of the formation. age. selandian. correlation. the våle formation is equivalent to the lellinge greensand and the kerteminde marl onshore denmark and lithologically most closely resembles the latter. the oldest part of the kerteminde marl and the lellinge greensand are coeval, but the latter has a more restricted distribution (sjælland and storebælt regions only, fig. 1; thomsen 1994; clemmensen & thomsen 2005). alisocysta reticulata is consistently present in the lowest part of the kerteminde marl (clemmensen & thomsen 2005). the ho of a. reticulata is therefore an important intravåle as well as intra-kerteminde marl marker that may be used to correlate the two formations. the våle formation correlates with the marly facies of the maureen formation in the uk and norwegian sectors of the central and viking grabens (knox & holloway 1992). 2900 2800 2940.5 2903.0 2894.4 2913.3 2963.2 horda fm balder fm sele fm lista fm våle fm bor mb vile mb ve mb bue mb chalk gp m augusta-1 gr sonic neutron/densityfig. 17. augusta-1, type well for the bor, ve and bue members. black bars show cored sections. 30 våle fm chalk group vile mb bor mb lista fm augusta-1 clay si. vf. f. m. sand c. vc. c c c 45° 27° 2920 2930 2940 2950 2960 2920 2930 2940 2950 2960 log depth core depthgr sonic bor member new member history. the bor member encompasses sandstone bodies enveloped in the marlstones of the våle formation in the danish north sea sector. hardt et al. (1989) recognised a pure sandstone unit, the ty formation, located between the ekofisk and lista formations in the southern viking graben. the ty formation replaces the våle formation in its occurrence area and may be contemporaneous with the bor member, but it is not contiguous with it and it has a different source area. the presence of sandstone bodies in the våle formation in the danish sector was recognised by a stratigraphic working group at statoil norway in the mid-1990s and the sandstones were informally named the ‘borr member’. derivation of name. after bor (danish spelling), the father of odin. type well. danish sector well augusta-1, 2963.2–2940.5 m mdrt (figs 17, 18). reference well. danish sector well cecilie-1, 2319.6–2284.6 m mdrt (fig. 19). distribution and thickness. the bor member has been encountered at the mouth of the siri canyon as well as in fig. 18. core log of bor member sandstones in the augusta-1 well. legend in fig. 9. the interval around the våle–lista formation boundary marked by a grey bar in the core depth column is shown as core photographs in fig. 23. 31 the nearby wells tabita-1, augusta-1 and cleo-1 (fig. 20a; plate 1). it reaches a thickness of up to 23 m. lithology. the bor member consists of olive-green, partly calcite-cemented sandstones. the sandstones are very fine grained to fine grained and well sorted (fig. 18). rounded and translucent quartz grains dominate, but the content of glaucony grains is high (20–25%). mica and pyrite concretions are present in small amounts. angular chalk and claystone clasts occur locally. although composed exclusively of sandstone in the type well (fig. 18), the member may also include subordinate interbedded marlstones (e.g. cecilie-1, fig. 19). log characteristics. the bor member sandstones are best identified on the density log where they produce a blocky pattern with density values significantly lower than those of the marlstones beneath and above. the sandstones may also be identified from a combination of the density and neutron logs, as the presence of pure sandstone results in a ‘cross-over’ of the two log curves (figs 17, 19). on the gamma-ray log,thebormember ischaracterised by a blocky log signature with only small-scale increasing or decreasing trends and with values clearly higher than those of the subjacent, suprajacentand locally interbeddedmarlstones. boundaries. the boundaries with the marlstones of the våle formation, the chalks of the ekofisk formation and the mudstones of the lista formation are sharp and characterised by prominent shifts on the gamma-ray, sonic and density logs (figs 17–19, 21). depositional environment. the sandstones of the bor member were deposited from highly concentrated gravity flows at bathyal depths. age. selandian. correlation. the bor member is contemporaneous with parts of the kerteminde marl onshore denmark. the lellinge greensand, which appears between the top chalk surface and the kerteminde marl in some areas in eastern denmark, may also be broadly contemporaneous with the bor member, but differs from it lithologically in being predominantly a glaucony-rich calcilutite, rich in bryozoan fragments. the bor member may be compared with the ty formation (hardt et al. 1989) and with sandstones in the maureen formation (deegan & scull 1977) in the norwegian and uk sectors of the southern viking graben and the central graben. however, it is not contiguous with these units and it has a different source area. 2100 2200 2300 bue mb ve mb vile mb tyr mb bor mb horda fm balder fm sele fm lista fm våle fm chalk gp m cecilie-1 gr sonic neutron/density 2241.7 2276.6 2284.6 2319.6 fig. 19. cecilie-1, reference well for the bor and tyr members. black bar shows cored section. 32 fig. 20. location map showing the distribution of the rogaland group sandstones in the siri canyon (the outline of the canyon is indicated by grey shading, the grey shading inside the canyon indicates an area of positive relief within the canyon). a: bor and tyr members. b: idun member. c: rind and kolga members. cleo-1 francisca-1 frida-1 d-1 nolde-1 siri-3siri-2 connie-1 siri-1 augusta-1 elna-1 tabita-1 amalie-1 cecilie-1 10 km augusta-1 tabita-1 nini-3 amalie-1 cleo-1 elna-1 frida-1 d-1 nolde-1 cecilie-1 siri-3 sandra-1 sandra-1 siri-2 siri-1 connie-1 10 km augusta-1 tabita-1 amalie-1 cleo-1 elna-1 frida-1 d-1 nolde-1 cecilie-1 siri-3 sandra-1 siri-2 siri-1 connie-1 10 km tyr mb bor mb both mbs idun mb rind mb kolga mb both mbs sir i c an yo n sir i c an yo n sir i c an yo n a b c nini-3 francisca-1 francisca-1 nini-3 lista formation history. deegan & scull (1977) established the lista formation for the widespread, non-laminated mudstones that overlie the marls of the unit equivalent to the maureen formation (våle formation). kristoffersen & bang (1982) established the non-calcareous clay and shale unit cen-1 between the top of their north sea marl (våle formation) and the base of the beds with volcanic tuff. they noted that the cen-1 unit corresponds to the lista formation. for reasons of seniority and the informal nature of the cen units, we maintain the name lista formation for this stratigraphic unit. type well. norwegian sector well 2/7–1, 2917.5–2872.5 m mdkb. danish reference wells. e-8, 2057.0–2027.6 m mdkb (fig. 13); cleo-1, 2812.0–2765.5 m mdkb (fig. 21; plate 1). distribution and thickness. the lista formation is present throughout the north sea basin, except in a few areas where it has been removed by erosion. in the danish sector, its thickness varies from 0 to 108 m (fig. 22). lithology. the formation is characterised by dark coloured, predominantly greyish, greenish or brownish, non-laminated to faintly laminated, non-calcareous mudstones. the lista formation is predominantly non-tuffaceous but becomes tuffaceous towards its top. in the siri canyon, glaucony-rich, massive sandstone layers and injected sandstone bodies occur in the lista formation. log characteristics. although fluctuating, both the gamma-ray and sonic log readings in the lista formation have higher mean values than those of the underlying våle formation and lower mean values than those of the overlying sele formation. in wells where mudstone facies dominate in the lista formation, the gamma-ray and sonic log patterns can be subdivided into three. the tripartite log pattern reflects the succession of three different mudstone units, established as new members herein (see below). boundaries. in most wells where the transition has been cored, the boundary is sharp between the light-coloured marlstones of the våle formation and the dark-coloured, non-calcareous mudstones of the lower lista formation (vile member, see below; fig. 23). on the gamma-ray log, the boundary is picked at an abrupt upward shift to higher values than in the underlying våle formation. this level can typically be identifiedon the sonic log at a velocity 33 minimum. above this minimum, the sonic readings increase slightly upwards. the lista formation is overlain by the sele formation. the base of the sele formation was defined by deegan & scull (1977 p. 34) at the contact between “non-laminated, non-tuffaceous shales” (lista formation) and “laminated tuffaceous shales” (sele formation). this boundary definition was followed by mudge & copestake (1992a, b). on the other hand, knox & holloway (1992 p. 46) followed o’connor &walker (1993) and placed the boundary somewhat lower, at the contact between “grey-green and green-grey, blocky, bioturbated claystones” of the lista formation and “dark grey fissile mudstones” of the sele formation. the boundary concept of knox & holloway implies that the “non-laminated, non-tuffaceous shales” of deegan & scull are incorporated in the sele formation where these, together with overlying laminated indisputable sele mudstones, constitute the basal sele unit s1a (knox & holloway 1992). in the present paper, the boundary concept of deegan & scull (1977) is followed, and the lower part of the unit of “non-laminated, nontuffaceous shales” (the “dark grey, fissile mudstones” of knox & holloway) that overlies the grey-green mudstones is retained in the lista formation as its topmost unit. this unit is formalised as a new member of lista formation herein (bue member, see below). subdivision. the lista formation is subdivided into six new members. three of these, the vile, ve and bue members, are mudstone units that have widespread distribution in the north sea basin and can be correlated with danish onshore units. in the siri canyon, fine-grained fig. 22. isochore map of the lista formation in the study area. the positions of the two danish reference wells, cleo-1 and e-8, are indicated. 2700 2800 bue mb ve mb vile mb bor mb horda fm balder fm sele fm lista fm våle fm chalk gp m cleo-1 2765.5 2812.0 2792.4 2777.6 gr sonic densityfig. 21. cleo-1, danish reference well for the lista formation and the vile, ve and bue members. lista formation cleo-1 e-8 20 40 60 80 100 thickness (m) 25 km 34 2918 m 2919 m augusta-1 2920 m 2921 m 0 10 20 30 40 50 60 70 80 90 100 cm sandstone bodies occur within each of the three mudstone units. these sandstone-dominated units are proposed here as three new members: the tyr member for the sandstones in the vile member, the idun member for the sandstones in the ve member, and the rind member for the sandstones in the bue member. knox & holloway (1992) recognised a threefold subdivision of the lista formation exclusively based on biostratigraphy. their l1 and l2 units are separated by the ho of the dinoflagellate palaeoperidinium pyrophorum, and the l2 and l3 units are separated by the ho of areoligera gippingense. it is noticeable that these two bioevents occur close to the boundaries between the three mudstone members of the lista formation proposed herein on the basis of lithology. macroand ichnofossils. macrofossils have not been reported fromthelistaformation;theformationismoderatelytoheavily bioturbated(for ichnotaxa, see individualmembersbelow). microfossils and palynomorphs. the lista formation differs from the underlying våle formation by lacking common planktonic foraminifers and from the overlying sele formation by having an impoverished agglutinated benthic foraminifer assemblage. the lista formation contains a characteristic sequence of palynomorph datums that can aid separation of its members. these datums are treated under the individual lista members below. depositional environment. the lista formation consists predominantly of hemipelagic mudstones and was probably deposited from very dilute turbidity currents and from suspension. the composition of the microfaunal assemblage indicates a relatively open marine depositional setting in upper to possibly middle bathyal depths with oxic to dysoxic bottom conditions. this is based on the presence of an impoverished agglutinated foraminifer assemblage dominated by tubular suspension feeders (especially rhabdammina spp.) together with epifaunal and infaunal detritivores (e.g. haplophragmoides spp. and spiroplectammina spectabilis). the relative abundance of tubular suspension feeders is higher in wells in the siri canyon than in wells outside the canyon. this probably indicates slightly deeper water within the canyon area during deposition of the lista formation. age. selandian–thanetian (upper paleocene). the selandian–thanetian boundary may be placed in the middle part of the lista formation (in the lower part of the ve member, see below), at the ho of the dinoflagellate palaeoperidinium pyrophorum. correlation. the lista formation corresponds to the following succession of upper paleocene units from onshore fig. 23. core photographs of the våle–lista formation boundary interval in the augusta-1 well. the boundary is at 2919.46 m where dark grey non-calcareous mudstones of the vile member overlie greenish grey marls of the våle formation. in most cores the boundary is sharp, as illustrated here. depths are core depths. stratigraphic position of the figured interval is shown in fig. 18. 35 fig. 25. correlation diagram showing possible diachronism of the våle–vile boundary in an east–west transect extending into the norwegian sector of the north sea (1/3-1 and 2/7-1 are norwegian sector wells). the distribution of biostratigraphic events shows that the våle–vile boundary youngs in a westerly direction (ho, highest occurrence). alternatively, the event distribution could be explained as a result of reworking of older strata into the lista formation in the danish sector. the figure also shows an example of a well (2/7-1) with a relatively large separation between the base of the sele formation and the lowest and most conspicuous gamma-ray peak in the formation (see text for further explanation). æbelø fm holmehus fm østerrende clay ølst fm kerteminde marl danian limestone 1/3-1 gr sonic gr sonic gr sonic gr sonic 1/3-1 2/7-1 mona-1 e-8 viborg-1 viborg-1 sele fm lista fm bue mb ve mb vile mb våle fm chalk group ho abundant palaeoperidinium pyrophorum ho isabelidinium? viborgense ho diverse calcareous benthic foraminifers ho (provisional) planktonic foraminifers 2050 2075 450 400 3050 3000 2900 2950 3150 3200 3250 kolga mb s3 s2b s2a s1b bue mb ve mb vile mb horda fm balder fm sele fm lista fm våle fm chalk gp 2000 2100 m siri-3 1998.8 2016.8 2036.1 2066.4 2072.8 2102.6 gr sonic densityfig. 24. siri-3, type well for the vile and kolga members, and danish reference well for the balder formation. the figure also shows the subdivision of the sele formation used by knox & holloway (1992); in this well, the bue member is equivalent to the s1a subunit of these authors. black bar shows cored section. 36 denmark: æbelø formation (informal mudstone unit described by bøggild 1918 and heilmann-clausen 1995), holmehus formation (heilmann-clausen et al. 1985) and østerrende clay (informal mudstone unit described by nielsen et al. 1986 and heilmann-clausen 1995). vile member new member history. the vile member comprises the widespread, dark olive-grey to dark grey, non-calcareous, fissile mudstones that constitute the lower part of the lista formation. the unit was recognised by a stratigraphic working group at statoil norway in the mid-1990s and informally named the ‘vile formation’. derivation of name. after vile, the brother of odin. type well. danish sector well siri-3, 2102.6–2072.8 m mbrt (fig. 24; plate 4). reference wells. danish sector wells e-8, 2057.0–2044.0 m mdkb (fig. 13); cleo-1, 2812.0–2792.4 m mdkb (fig. 21; plate 1). distribution and thickness. the vile member has been recognised in a large number of north sea wells, and the unit probably has a basinwide distribution. however, it is apparently lacking in the siri canyon wells connie-1 and siri-2 (figs 29, 31), probably due to erosion. its thickness varies between 0 and 30 m over most of the danish sector. it greatest thickness is reached in the siri canyon. lithology. the member consists of dark olive-grey to dark grey, non-calcareous, swelling, smectitic, fissile mudstones (fig. 23). thin silicified layers occur in the member. calcite is common and occurs as small nodules and larger concretions. in the siri canyon, small pyrite concretions and less than 1 cm thick, silty, very fine-grained glaucony-rich sandor siltstone laminae are locally present in the vile member. the laminae are parallel to the bedding of the mudstones; they have sharp bases and are normally grad1803.7 1763.9 1700.4 1717.2 horda fm lark fm balder fm sele fm kolga mb bue mb vile mb idun mb tyr mb lista fm chalk gp 1600 1700 1800 m nini-3 gr sonic neutron/density s3 s2b s2a s1b fig. 26. nini-3, type well for the tyr member and reference well for the kolga member. the figure shows the subdivision of the sele formation used by knox & holloway (1992); in this well, the bue member is equivalent to the s1a subunit of these authors. black bar shows cored section. 37 ed. thin concordant or discordant, postdepositional sandstone intrusions are locally present. in the lower part of the vile member, the intrusions are only a few millimetres thick, but they often increase in number and thickness towards the top of the member (fig. 11). log characteristics. in most wells there is a gradual increase in gamma-ray response up through the vile member, accompanied by a slight decrease in sonic readings. boundaries. the lower boundary of the vile member is that of the lista formation. the upper boundary is defined by the base of the ve member; boundaries with the sandstone-dominated tyr member are described under that member. macroand ichnofossils. the vile member is moderately to intensely bioturbated. ichnogenera in the member include chondrites ispp., phycosiphon ispp., planolites ispp. and zoophycos ispp. microfossils and palynomorphs. the vile member is characterised by a general decrease in the diversity of benthic foraminifers and radiolaria from its base to its top. in the danish sector, the transition from the underlying våle formation to the vile member is marked by the provisional ho of planktonic foraminifers. a conspicuous drop in diversity of benthic foraminifers takes place in the middle of the vile member. the ho of the dinoflagellate isabelidinium? viborgense is an important intra-vile marker located in the upper part of the member. above it, a sudden decrease in the abundance of radiolaria further charlista fm chalk group nini-3 clay si. vf. f. m. sand c. vc. p p p 1800 1790 1780 1770 1760 1800 1790 1780 1770 1760 log depth core depthgr sonic vile mb tyr mb fig. 27. core log of the tyr member in the nini-3 well. for legend, see fig. 9. 38 2908 m2907 m 2909 m augusta-1 2910 m 2911 m 2912 m 0 10 20 30 40 50 60 70 80 90 100 cm acterises a level within the uppermost part of the vile member. the transition from the vile member to the overlying ve member is marked by a conspicuous drop in the abundance of the dinoflagellate palaeoperidinium pyrophorum(whichhas itshoataslightlyhigherstratigraphic level, within the lower part of the ve member, see below). depositional environment. the mudstones of the vile member are hemipelagic deposits, whereas the thin sandstone and siltstone laminae are interpreted as the deposits of low-density turbidity currents. the presence of zoophycos ispp. suggests depositional water depths of at least 200 m (bottjer & droser 1992). age. selandian. correlation.the vile member corresponds to theæbelø formation, onshore denmark (informal mudstone unit describedbybøggild1918andheilmann-clausen1995). it corresponds to the lista l1 subunit of knox & holloway (1992). fig. 28. core photographs of mudstones of the vile and ve members in the augusta-1 well. depths are core depths. the boundary between the two members is placed where greenish and reddish grey mudstones become dominant, at 2910.4 m (arrow). this depth corresponds to log depth 2913.3 m on fig. 17. 39 biostratigraphic correlation with well sections in the norwegian north sea sector may indicate that the lower boundaryof thevilemember isdiachronous (fig. 25). in the type well for the lista formation (norwegian well 2/7-1; fig. 25), its base (i.e. its contact with the marlstones of the underlying våle formation) is above the ho of a diverse calcareous benthic foraminifer assemblage. in wells in the danish sector, this event occurs within the vile member. similarly, the provisional ho of planktonic foraminifers, an event that is close to the boundary between the våle and lista formations in denmark, is found well within the våle formation in well 2/7-1. in the type well for the våle formation (norwegian well 1/3-1; fig. 25), the ho of i.? viborgense coincides with the våle–lista boundary (i. prince, unpublished biostratigraphic data). this event occurs above the våle formation in the danish sector, in the middle to upper part of the vile member. in the danish onshore well viborg-1, the latter event and the ho of the diverse calcareous benthic foraminifer assemblage occur above the kerteminde marl in the upper part of the æbelø formation, a correlative of the vile member (fig. 25; heilmann-clausen 1985). the distribution pattern of the biostratigraphic events indicates that sedimentation of marls continued in the norwegian sector some time after marl sedimentation was replaced by sedimentation of non-calcareous mudstones in the danish sector. alternatively, calcareous foraminifer assemblages and associated lithologies have been reworked into higher levels of the våle formation or even into the lista formation in the danish sector. tyr member new member history. the tyr member consists of glaucony-rich, sandstone-dominated deposits that are laterally equivalent to, and commonly underlain and overlain by, mudstones of the vile member. these sandstones were previously recognised by a stratigraphic working group at statoil norway in the mid-1990s and informally referred to the ty formation of hardt et al. (1989). derivation of name. after tyr, the son of odin. type well. danish sector well nini-3, 1803.7–1763.9 m mdrt (figs 26, 27; plate 4). reference well. danish sector well cecilie-1, 2276.6–2241.7 m mdrt (figs 8, 19). distribution and thickness. the tyr member has only been encountered in the siri canyon and it may be restricted to thatarea. it reaches a thickness of up to 40 m (fig. 20a). lithology. the tyr member is characterised by thick beds of olive-green to greenish grey, very fine-grained to finegrained and well-sorted sandstone (fig. 27). rounded and translucent quartz grains dominate, but the content of glaucony grains is high (15–20%), hence the greenish colour of the sandstones. mica and small pyrite concretions are present in small amounts throughout the member. angular chalk and claystone clasts occur locally in the sandstones. the sandstones are partly calcite-cemented. intrusive sandstones are common, particularly towards the top of the member where they may be several metres thick (fig. 8). subordinate interbedded dark grey noncalcareous mudstones resemble those of the laterally equivalent vile member. log characteristics. the tyr member is best identified on the density log where the sandstones are characterised by a conspicuously lower density than the associated mudstones. the sandstones may also be identified from a combination of the density and neutron logs, as the presence of pure sandstone results in a ‘cross-over’ of the two log curves (figs 19, 26). the gamma-ray response resembles that of the underlying våle formation, but is slightly lower than the response of the vile member (figs 8, 26). this log pattern makes it feasible to differentiate even minor sandunits from mudstone beds in the tyr member. thicker sand units may show decreasingor increasingupwards gamma-ray values. these trends do not seem to berelatedtograin-sizevariations, judgingfromcore studies. boundaries. the boundaries to the mudstones of the vile member, the marlstones of the våle formation and the chalks of the ekofisk formation are sharp and characterised by prominent shifts in gamma, sonic and density log readings (figs 19, 26, 27). in some wells, the tyr member overlies the våle formation or the ekofisk formation with an erosional contact (e.g. nini-3; figs 26, 27). depositional environment. although the sandstones of the tyr member were deposited from highly concentrated gravity flows, their present appearance is dominated by theeffectsofpostdepositional liquefaction and fluidisation. age. selandian. correlation. the tyr member is contemporaneous with parts of the lithologically dissimilar æbelø formation in 40 onshore denmark and with the lower part of the heimdal formation of deegan & scull (1977) as well as the andrew sandstone and the mey sandstone member of knox & holloway (1992) in the norwegian and uk sectors of the southern viking graben. however, it is not contiguous with the latter three sandstone units and has a different source area. ve member new member history. the ve member consists of variegated mudstones that have previously been recognised from north sea wells as the holmehus formation by heilmann-clausen et al. (1985), who gave no further details, and by danielsen & thomsen (1997), who indicated its presence in several wells. the unit was also recognised by a stratigraphic working group at statoil norway in the mid-1990s and informally named the ‘ve formation’. derivation of name. after ve, the brother of odin. type well. danish sector well augusta-1, 2913.3–2903.0 m mdrt (fig. 17). reference wells. danish sector wells e-8, 2044.0–2030.3 m mdkb (fig. 13); cleo-1, 2792.4–2777.6 m mdkb (fig. 21; plate 1). distribution and thickness. the sediments of the ve member have been recognised from a large number of north sea wells, and the unit probably has an almost basinwide distribution. its thickness varies from 0 to 21 m in the danish sector. lithology. the ve member consists of mottled green, bluish green, reddish brown and brown mudstones (fig. 28). mottled, purple coloured intervals are also present locally. the middle part of the member is often characterised by a thick dark reddish brown to chocolate brown interval. thevemembermudstones are non-calcareous and richin smectite. pyrite and carbonate concretions occur throughout the member. a weak biogenic lamination is sometimes observed in cores. only very little organic material is present in the member. in the siri canyon, thin intrusive sandstones are common in the ve member. log characteristics. in general, the gamma-ray log shows a decreasing-upwards trend through the ve member, as opposed to the increasing trend through the underlying vile member. in the uppermost part of the ve member, the gamma-ray response increases over a short interval before reaching the base of the overlying bue member. the sonic log pattern throughout the ve member is smooth and relatively stable compared with the sonic pattern of the vile member. it also differs from the latter in having an increasing-upwards trend. the log pattern of the ve member differs from that of the overlying bue member in having a lower gamma-ray response level. boundaries. the lower boundary, with the vile member, is placed at the first appearance of greenish, bluish or reddish brown mudstones above the dark olive-grey to dark grey mudstones of the vile member. the colour change from vile to ve mudstones is often gradational and the boundary may be difficult to define precisely (fig. 28), especially when only cuttings samples are available. however, in the colour transition interval, a gamma-ray spike separates an interval with an increasing-upwards gammaray trend below from an interval with a decreasing gamma-ray trend above (compare figs 17 and 28). in the absence of a clear indication of the boundary level from sediment colour change, the gamma spike at the shift from increasing to decreasing gamma-ray values may be used as a marker for the boundary. the vile member is absent from the connie-1 well, and in this well the lower contact of the ve member is with the marlstones of the våle formation (fig. 31). the upper boundary is at the base of the bue member (see below). boundaries with the idun member are described under that member. macroand ichnofossils. the mudstones of the ve member are normally heavily bioturbated. the most common trace fossils are phycosiphon ispp. and zoophycos ispp. chondrites ispp. and planolites ispp. are present, but rare. microfossils and palynomorphs. in the danish sector of the north sea, the ho of abundant palaeoperidinium pyrophorum is located at or close to the vile–ve boundary. the hos of p. pyrophorum and palaeocystodinium australinum are in the lower part of the ve member. in general, the dinoflagellate assemblage from the upper part of the ve member is sparse and is characterised by specimens of areoligera gippingensis. an acme of the latter species marks a level in the upper part of the ve member. the highest in situ occurrence of the dinoflagellate alisocysta margarita is located close to the top of the ve member. depositional environment. deposition of the mudstones of the ve member was controlled by hemipelagic sedimentation and sedimentation from dilute turbidites. the 41 idun member new member history. the idun member consists of sandstone-dominated deposits that are laterally equivalent to, and commonly underlain by, mudstones of the ve member. this sandstone unit was previously recognised by a stratigraphic working group at statoil norway in the mid-1990s and was informally referred to the heimdal formation of deegan & scull (1977). derivation of name. after idun, the goddess of youth. type well. danish sector well connie-1, 2368.3–2332.0 m mdrt (figs 29, 30). 2200 2300 2368.3 2332.0 2329.6 2292.2 idun mb rind mb bue mb bue mb ve mb horda fm balder fm sele fm lista fm våle fm chalk gp m connie-1 gr sonic neutron/densityfig. 29. connie-1, type well for the idun and rind members. in this well, the rind member may be divided into three major sandstone intervals. the idun member consists of two thick sandstone intervals, separated by a thick mudstone unit. black bars show cored sections. occurrence of the trace fossil zoophycos ispp. indicates a water depth of at least 200 m (bottjer & droser 1992). the overall high degree of bioturbation, the lack of organic material and the greenish, bluish and reddish brown colours together suggest oxygenated bottom conditions. age.selandian–thanetian.theselandian–thanetianboundaryisplacedatthehoofthedinoflagellatepalaeoperidinium pyrophorum, in the lower part of theve member. correlation. the ve member correlates with the holmehus formation (heilmann-clausen et al. 1985) onshore denmark and is lithologically indistinguishable from that formation. 42 lista fm våle fm connie-1 clay si. vf. f. m. sand c. vc. 45o s v g g g pp ppp 45o 2290 2300 2310 2320 2340 2350 2360 2300 2290 2310 2320 2330 2340 2350 2360 2370 log depth core depthgr sonic p p p p p p p p p p p p change of core depth scale bue mb idun mb bue mb rind mb ve mb fig. 30. core log of the idun and rind members in the connie-1 well. for legend, see fig. 9. minor mudstone beds separate the three major sandstone intervals of the rind member. minor mudstone layers are also intercalated with the idun member sandstone intervals. the core depth scale of the lower core is offset by c. 1.6 m relative to the scale of the upper core in the figure. for reasons of consistency with core data from this well, the original (albeit erroneous) core depths of the lower core are maintained in the figure. this does not affect the depths of the top and base of the idun member given in the text, as these are based on log depths. 43 reference well. danish sector well siri-2, 2205.5–2127.0 m mdrt (fig. 31). distribution and thickness. the idun member is only known from the siri canyon, and it may be restricted to that area (fig. 20b). it reaches a thickness of up to 179 m in the siri-2 well. lithology. the idun member is dominated by very finegrained to fine-grained, well-sorted sandstones (fig. 30). rounded and translucent quartz grains dominate, but the content of glaucony grains is high (15–25%). the sandstones are olive green to greenish grey due to the high content of glaucony. mica and small pyrite concretions are present in small amounts. angular chalk and claystone clasts occur locally. intrusive sandstones are also represented (figs 11, 12, 30). the sandstones and adjacent mudstones are partly calcite-cemented. in the nini and siri wells, the sandstones occur in thick amalgamated successions with only rare, thin mudstone interbeds; the latter are lithologically comparable to the laterally equivalent ve member mudstones (see above). in the connie-1 well, however, the sand-rich succession is interrupted by a discrete 6 m thick mudstone unit (fig. 30). log characteristics. the sandstone-dominated idun member is best identified on the density log where it is characterised by a conspicuously lower density than the underlying and overlying mudstones (figs 29, 31). the sandstone component may also be identified from a combination of the density and neutron logs, as the presence of pure sandstones results in a ‘cross-over’ of the two log curves (figs 29, 31). the idun member is characterised by a blocky, decreasing-upwards gamma-ray and density log pattern. intervals with an overall constant gamma-ray pattern may be characterised by many small-scale increasingor decreasing-upwards gamma-ray cycles. boundaries. in sections where the sandstones of the idun member are enveloped by mudstones of the ve member, the boundaries are sharp and characterised by prominent shifts on the gamma-ray, sonic and density logs (figs 29, 30). where the ve member is absent, comparable, sharp boundaries are observed with the mudstones of the vile member beneath and the bue member above (figs 11, 26, 30). in the siri-2 well, the lower lista formation is absent and an erosive unconformity separates the idun member sandstones from the marlstones of the våle formation (fig. 31). depositional environment. although the sandstones of the idun member were deposited from highly concentrated gravity flows, their present appearance largely records postdepositional liquefaction and fluidisation processes. age. selandian–thanetian. bue mb bue mb rind mb idun mb balder fm horda fm sele fm lista fm chalk gp våle fm 2100 2200 m siri-2 gr sonic neutron/density 2127.0 2205.5 fig. 31. siri-2, reference well for the idun member. black bar shows cored section. 44 with these sandstone units, however, and has a different source area. bue member new member history. the bue member encompasses the light to dark grey and greyish black mudstones that occur between the top of the ve member and the base of the sele formation. these mudstones have not previously been recognised as a separate unit in the danish sector. derivation of name. after bue, the son of odin and rind. typewell. danishsectorwellaugusta-1,2903.0–2894.4m mdrt (fig. 17). referencewells.danishsector wells e-8, 2030.3–2027.6 m below mdkb (fig. 13); cleo-1, 2777.6–2765.5 m mdkb (fig. 21; plate 1). distribution and thickness. the sediments of the bue member have been recognised from a largenumberof north sea wells, and the unit probably has a basinwide distribution. its thickness varies from 0 to 18 m in the danish sector. lithology. the bue member consists of light to dark grey and greyish black mudstones. the mudstones are generally rich in smectite. in the siri canyon, the upper part of the member sometimes contains laminae of very finegrained to fine-grained sandstones or siltstones, mimicking the laminated mudstones of the overlying sele formation (figs 32, 39). the laminae are less than 2 cm thick, have sharp bases, are normally graded and show parallel lamination. concordant or discordant sandstone intrusions are locally present in the member (fig. 8). small calcite and siderite concretions are occasionally present. moderately to intensely bioturbated intervals are interbedded with non-bioturbated intervals. tuff layers may be present in the member. log characteristics. the gamma-ray response of the bue member is generally higher than that of the underlying ve member, but lower than that of the overlying sele formation. in some siri canyon wells, minor coarseningupwards cycles are indicated by the gamma-ray log of the bue member. boundaries. the transition from typical lithologies of the ve member to those of the bue member is often gradafig. 32. core photographs showing the bue member mudstones with numerous sandstone laminae in the augusta-1 well. depths are core depths. 0 10 20 30 40 50 60 70 80 90 100 cm 2891 m 2892 m augusta-1 2893 m correlation. the idun member is contemporaneous with parts of the lithologically dissimilar holmehus formation onshore denmark and with parts of the heimdal formation of hardt et al. (1989) and the lower balmoral sandstone and tuffite of the mey sandstone member of knox & holloway (1992). the idun member is not contiguous 45 rind member new member history. the rind member consists of sandstone-dominated deposits that are laterally equivalent to, and commonly underlain and overlain by, mudstones of the bue member. sandstone bodies at this stratigraphic level were previously recognised by a stratigraphic working group at statoil norway in the mid-1990s and were informally referredtotheheimdalformationofdeegan&scull (1977). derivation of name. after the giantess rind. type well. danish sector well connie-1, 2329.6–2292.2 m mdrt (figs 29, 30). reference well. danish sector well sandra-1, 2066.3– 2004.8 m mdrt (fig. 33). distributionand thickness. the rind member has only been encountered in the siri canyon, and it may be restricted to that area where it reaches a thickness of 62 m (fig. 20c). lithology. the rind member consists of very fine-grained, well-sorted sandstones interbedded with thin mudstone beds that typically form less than 15% of the member (fig. 30). rounded and translucent quartz grains dominate in the sandstones, but the content of glaucony grains in the very fine-grained to fine-grained size fraction is high (15–25%). the sandstones are olive green to greenish grey due to the high content of glaucony. mica and small pyrite concretions are present in small amounts. angular chalk and claystone clasts occur locally in the sandstones, which are partly calcite-cemented. the interbedded mudstones are lithologically comparable to the bue member mudstones (see above). log characteristics. the rind member is best recognised on the density log where it shows either a blocky or a serrate pattern created by the alternation of sandstone beds or amalgamated units (low density) with thin mudstone beds (high density; figs 29, 33). the sandstones may also be identified from a combination of the density and neutron logs, since the presence of pure sandstones results in a ‘cross-over’ of the two log curves (figs 29, 33). the gamma-ray log shows a low-amplitude serrate pattern. this pattern does not reflect alternating sand or mudstones, judging from core inspection. boundaries. the boundary between the sandstones of the rind member and the mudstones of the bue member is tional and the boundary may therefore be difficult to position precisely. it is placed where mottled green, bluish green, reddish brown and brown mudstones pass upwards into grey mudstones with sandstone and siltstone laminae. on the petrophysical logs, this transition is reflected by a shift from decreasingto increasing-upwards gamma-ray values or at an abrupt increase in the gamma-ray response. the upper boundary of the bue member is at the base of the sele formation. macroand ichnofossils. trace fossils recognised in the bue member are phycosiphon ispp., planolites ispp., thalassinoides ispp. and rare zoophycos ispp. microfossils and palynomorphs. the ve–bue boundary is bracketed by the stratigraphic succession of the ho of in situ alisocysta margarita (occurring in the upper ve member) followed by the ho of an impoverished assemblage of benthic agglutinated foraminifers (in the lower part of the bue member). the upper part of the bue member is characterised by common spores and pollen, in particular bisaccate pollen and inaperturopollenites spp. the bue– sele boundary is marked by the base of an acme of the dinoflagellate genus apectodinium and the lo of apectodinium augustum. depositional environment. the normally graded sandstone to siltstone laminae in the upper part of the member indicate that deposition of the bue member took place from dilute, low-density, turbidity currents in a generally sediment-starved environment at this level. the minor coarsening-upwards cycles observed on petrophysical logs from some siri canyon wells probably indicate either small distal lobes of deep-water channel-sandstones or levee deposits. age. thanetian. correlation. the bue member corresponds to the østerrende clay (informal mudstone unit described by nielsen et al. 1986 and heilmann-clausen 1995) onshore denmark. the level here defined as the boundary between the ve and thebuememberswascorrelatedbyknox (1997 fig. 3; the lista–sele boundary of this worker) with the boundary between the holmehus formation and the østerrende clay (as ‘grey clay’) onshore denmark. the bue member further correlates with the lower part of the s1a subunit of the sele formation established by knox & holloway (1992; see correlation section under the lista formation for further details). 46 sharp and characterised by prominent shifts on both the sonic and density logs (figs 29, 33). it is often difficult to identify the boundaries on the gamma-ray log alone. depositional environment. although the sandstones of the rind member were deposited from highly concentrated gravity flows, their present appearance largely records postdepositional liquefaction and fluidisation processes. age. thanetian. correlation. the rind member may be contemporaneous with parts of the lithologically dissimilar østerrende clay encountered in the storebælt region (fig. 1), with sandstones in the higher parts of the heimdal formation (hardt et al. 1989) and with the upper balmoral sandstone of the mey sandstone member of knox & holloway(1992).however, therindmember is not contiguous with those sandstone units and has a different source area. sele formation history. the sele formation was established by deegan & scull (1977) for the dark grey to greenish grey, laminated and carbonaceous, tuffaceous, montmorillonite-rich shales and siltstones that overlie the non-laminated and nontuffaceous shales of the lista formation in some areas, or arenaceous sediments belonging to a variety of different units in other areas. the original definition of the sele boundary is followed herein. this implies that the base of the sele formation is located at the base of the “laminated tuffaceous shales” that overlie the “non-laminated, nontuffaceous shales” of the lista formation (deegan & scull 1977; see boundaries section under the lista formation for further details). sandstones occur in the sele formation in the danish sector; these are established as a new member, the kolga member. typewell.british sectorwell21/10-1,2131–2100m mdkb. bue mb rind mb vile mb ve mb balder fm horda fm lark fm sele fm lista fm våle fm 1800 1900 2000 2004.8 2066.3 2100 m sandra-1 gr sonic neutron/density fig. 33. sandra-1, reference well for the rind member. black bar shows cored section. geological survey of denmark and greenland bulletin 15, 2008, 77-80 hans ø – or tartupaluk to the indigenous population of north-west greenland – is a small steeply sided island in nares strait at c. 80°50´n. charted in 1871 and named after greenlander hans hendrik, it is one of five limestone islands forming an integral part of the greenland silurian succession. rising less than 170 m above normally ice-infested waters, the 1.25 km2 island is physiographically far oversha d owed by nearby franklin ø (fig. 1). the island’s notoriety results from its placing more or less equidistant between the coasts of kennedy channel on the political boundary between greenland and canada. for 40 years the rocky patch has been the subject of a dispute be tween the danish/greenland and canadian governments regarding sovereignty rights, an issue that remains unresolved. however, there is mutual understanding between ca nada and denmark that “since the question of sovereignty over the island has not yet been solved no action should be taken by either side which might prejudge the settlement of the issue” (brückner 1984). formally, this remains the position today. 2007 developments and this article in 2007, two geological map sheets entitled ‘hans island, nunavut’ were released by the canadian government. form ing part of a richly illustrated report, the maps with structural cross-sections portray the island’s geology and that of ken nedy channel at scales 1:5000 and 1:100 000, respectively (harrison et al. 2007). also released are offshore geophysical and bathymetric data, including seismic and re frac tion profiles, obtained with the canadian coast guard ice-breaker louis s. st-laurent. the new data are an important addition to scientific knowledge of nares strait. however, the section ‘history of geological research’ does not match the excellence of the rest of harrison et al.’s report. beginning with canadian geodetic activities in 1953, the section continues: “it has been re ported that robert l. christie of the geological survey of canada visited hans island between 1957 and 1966 during geological mapping on northeastern ellesmere island (see dawes, 2004)”. yet, dawes (2004) makes no mention of any such visit. indeed, this author – who worked very closely with the late dr. christie on the history of exploration of nares strait and participated in his field programme in 1965 and 1966 – doubts that he ever set foot on the island or even planned to. harrison et al. (2007) list further canadian activity on and around hans ø but limit comment on danish geological research to: “it is evident that danish geologists have also visited hans island. there is a bedding attitude recorded on the island on the map of dawes and garde (2004) who consider bedrock here to be an exposure of the cape morton formation of silurian age”. however, in a postscript, acknowledgement is given to lauge koch’s map from 1922 with reference to dawes & haller (1979). by any standards this is meagre reporting, particularly so since the island was initially mapped by danish geologists and has appeared as an integral part of the silurian of washington land on danish maps for more than 75 years (e.g. koch 1931; troelsen 1950; jepsen et al. 1983). the lack of reference to these studies seems even more odd when early canadian geological maps of north-eastern ellesmere island covering the latitude of hans ø – for example, christie (1964, 1967) – do not portray the island, and neither does the 1:250 000 map of kennedy channel that is part of the canadian national map sheet coverage (kerr 1973). 77 hans ø, celebrated island of nares strait between greenland and canada: from dog-sledge to satellite mapping peter r. dawes and tapani tukiainen © geus, 2008. geological survey of denmark and greenland bulletin 15, 77–80. available at: www.geus.dk/publications/bull greenland canada nares strait 70° 70° 70° 80° 81° 81° 80° 65° 65° kane basin humboldt gletscher hall basin washington land el lesmere is land 50 km innuitian orogen arctic platform crozier ø franklin ø hans ø joe ø offley ø fig. 1. geological map showing two main structural provinces of kennedy channel and five carbonate islands, with hans ø in mid-channel. this article is stimulated by the realisation that the early mapping of this part of the high arctic is apparently little known internationally. the aim is two-fold. firstly, to briefly summarise danish geological work at kennedy channel, illustrated by the first map featuring hans ø (fig. 2), and secondly, to illustrate how the geological survey in copen hagen employs modern research techniques to maintain its interest in the far north. 78 fig. 2. the first geological map showing hans ø. printed in 1931 and annotated for publication in volume 73, the map appeared in meddelelser om grønland volume 200 (dawes & haller 1979, plate 2; for explanation see text). from dawes (1984; copies in survey archives and the royal library, copenhagen, are 28 × 23 cm). first geological map of hans ø the initial danish mapping of northern greenland took place via nares strait on two dog-sledge expeditions: the 2nd thule expedition 1916–1918 and the bicentenary jubilee expedition 1920–1923. the expeditions carried strong political overtones and for the purpose of fund raising national pride was a timely motif. the overriding aim was regional mapping of the national territory of denmark, with the ambitious plan becoming the lot of geologist and cartographer lauge koch. the scientific results were outstanding. for example, a topographical series of 19 sheets at 1:300 000 and a geological series of five sheets at varying scale were printed: geology in two batches, in 1929 and 1931, and topography in 1932 (dawes & haller 1979, fig. 3). hans ø, coloured as part of greenland, appears on topographic sheet 11 ‘cape constitution’ and on geological sheet ‘washington land’ where it is referred to the offley island formation (figs 2, 3). the five geological sheets with map descriptions were in tended for publication in volume 73 of meddelelser om grøn land, and the maps were annotated accordingly. two were published as planned (koch 1929, 1933). the fate of the remaining maps is intimately bound up with the destruction of stockpiles in copenhagen during the german occupation in the 2nd world war destining the washington land map – compiled in 1922 and printed in 1931 – to be released years later in another volume of meddelelser om grønland (dawes & haller 1979). later danish geological work during the thule and ellesmere land expedition 1939– 1941, danish geologist johannes c. troelsen examined and refined koch’s (1929) lower palaeozoic lithostratigraphy in southern washington land. hans ø and the four other islands were included on his map in the same colour as the silurian limestone and shale of the greenland coast (troelsen 1950). air-supported regional mapping by the geological survey of greenland in 1975–1977 and 1984–1985 includ ed visits to hans ø and other islands, for example in 1975 and 1984, for the purpose of studying the bedrock with a view to correlation with the lithostratigraphic framework established on washington land and farther north (peel 1984). hans ø was referred to the cape morton formation of the washington land group and so portrayed on the 1:250 000 map sheet of jepsen et al. (1983). aster satellite data and 3-d modelling combined with traditional information from aerial photo graphs, satellite data form an important tool in geological and commodity mapping in greenland. recently, aster (advanced spaceborne thermal emission and reflection radiometer), the imaging instrument on nasa’s terra satellite launched in december 1999, has proved to be particularly attractive due to its availability and low cost. it provides high-resolution images in 14 different bands of the electromagnetic spectrum ranging from visible to reflected and thermal infrared light. the image resolution ranges between 15 and 90 m and can be used to create detailed maps of surface temperature of land, reflectance and elevation. the imaging power of aster data is illustrated by fig. 4. concluding remarks initial surveying of the now celebrated island of nares strait can be attributed to the inughuit with their specialised travel techniques since their historic name tartupaluk, meaning ‘kidney-shaped’, precisely describes the island’s coastal outline. danish geological research along nares strait spans 90 years from the same dog-sledging era through helicopter operations to satellite mapping. modern remote-sensing tech niques provide an accurate perspective of an island traditionally visited by indigenous folk and a handful of explorers and geoscientists, but now – as a disputed border – by politicians and military personnel. 79 fig. 3. extract of ‘cape constitution’, sheet 11 of the topographical map of north greenland, scale 1:300 000, showing hans ø. from koch (1932). the original size of the extract is 21 × 23 cm. references brückner, p. 1984: hans island, kennedy channel. journal of glaciology 30(105), p. 256 only. christie, r.l. 1964: geological reconnaissance of northeastern ellesmere island, district of franklin. geological survey of canada memoir 331, 79 pp. + map. christie, r.l. 1967: reconnaissance of the surficial geology of northeastern ellesmere island, arctic archipelago. geological survey of canada bulletin 138, 50 pp. + map. dawes, p.r. 1984: notes on hans ø, kennedy channel: discovery and naming, early geological maps and a 1983 fly-past, 14 pp. + maps. unpublished report, geological survey of greenland, copenhagen. dawes, p.r. 2004: explanatory notes to the geological map of green land, 1:500 000, humboldt gletscher, sheet 6. copenhagen: geo logical survey of denmark and greenland, 48 pp. + map. dawes, p.r. & garde, a.a. 2004: geological map of greenland, 1:500 000, humboldt gletscher, sheet 6. copenhagen: geological survey of den mark and greenland. dawes, p.r. & haller, j. 1979: historical aspects in the geological investigation of northern greenland. part 1: new maps and photographs from the 2nd thule expedition 1916–1918 and the bicentenary jubilee expedition 1920–1923. meddelelser om grønland 200(4), 38 pp. + 4 plates. harrison, j.c., dewing, k. & mayr, u. 2007: geology of hans island and adjacent parts of kennedy channel, northwest greenland (kalaallit nunaat) and northern nunavut (canada). geological survey of ca nada open file 5321, 33 pp. + maps. jepsen, h.f., henriksen, n., hurst, j.m. & peel, j.s. 1983: geology, 1:250 000 map, washington land and daugaard-jensen land. copenhagen: geo logical survey of greenland. kerr, j.w. 1973: geology, kennedy channel and lady franklin bay, district of franklin, map 1359a, 1:250 000. ottawa: geological survey of canada. koch, l. 1929: the geology of the south coast of washington land. meddelelser om grønland 73(1/1), 39 pp. + plates. koch, l. 1931: washington land. geological map at 1:500 000. copen hagen: geodetic institute. koch, l. 1932: map of north greenland, 1:300 000, 19 sheets. copen hagen: geodetic institute. koch, l. 1933: geology of inglefield land. meddelelser om grønland 73(1/2), 38 pp. + plates. peel, j.s. 1984: preliminary report. geological reconnaissance on hans ø, north greenland, 4 pp. + map. unpublished report, geological survey of greenland, copenhagen. troelsen, j.c. 1950: contributions to the geology of northwest green land, ellesmere island and axel heiberg island. meddelelser om grøn land 149(7), 85 pp. + map. 80 authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: prd@geus.dk hans ø fig. 4. perspective view of kennedy channel showing hans ø and franklin ø viewed from the south towards canada. the colour composite of the visible and near-infrared aster image data draped over the digital elevation model are extracted from four aster images taken between 26 june and 30 july 2003. inset: hans ø viewed from the south-east. geological survey of denmark and greenland bulletin 1, 61-73 61 middle jurassic sediments occur in almost all parts of what is conventionally drawn as the present-day continent of europe – from portugal to the caucasus, sicily to svalbard, the hebrides to the petshora – but in the context of the present book, we shall confine ourselves essentially to western and northern europe, broadly from the alps to the arctic (fig. 1). the shelf-seas of the barents shelf and svalbard, also still part of europe today, are not included, but the eastern shores of greenland are. the palaeolatitudes in the middle jurassic were about 15º lower than they are today, so that europe straddled the temperate zones from the sub-tropical to the sub-arctic. this is fully reflected both in the lithoand biofacies of the sediments, ranging from the predominantly warm-water carbonates in the south, with their immensely diverse fossil biotas, to the siliciclastics with their impoverished fossil assemblages in the north. historically, western europe, as the cradle of geological science together with its superbly developed jurassic successions, has given us a longstanding knowlthe middle jurassic of western and northern europe: its subdivisions, geochronology and correlations john h. callomon the palaeogeographic settings of denmark and east greenland during the middle jurassic are outlined. they lay in the widespread epicontinental seas that covered much of europe in the post-triassic transgression. it was a period of continuing eustatic sea-level rise, with only distant connections to world oceans: to the pacific, via the narrow viking straits between greenland and norway and hence the arctic boreal sea to the north; and to the subtropical tethys, via some 1200 km of shelf-seas to the south. the sedimentary history of the region was strongly influenced by two factors: tectonism and climate. two modes of tectonic movement governed basinal evolution: crustal extension leading to subsidence through rifting, such as in the viking and central grabens of the north sea; and subcrustal thermal upwelling, leading to domal uplift and the partition of marine basins through emergent physical barriers, as exemplified by the central north sea dome with its associated volcanics. the climatic gradient across the 30º of temperate latitude spanned by the european seas governed biotic diversity and biogeography, finding expression in rock-forming biogenic carbonates that dominate sediments in the south and give way to largely siliciclastic sediments in the north. geochronology of unrivalled finesse is provided by standard chronostratigraphy based on the biostratigraphy of ammonites. the middle jurassic saw the onset of considerable bioprovincial endemisms in these guide-fossils, making it necessary to construct parallel standard zonations for boreal, subboreal or nw european and submediterranean provinces, of which the nw european zonation provides the primary international standard. the current versions of these zonations are presented and reviewed. keywords: northwest europe, north sea, east greenland, middle jurassic, palaeogeography, geochronology, ammonite biostratigraphy, standard chronostratigraphy university college london, 20 gordon street, london wc1h 0aj, uk. e-mail: johncallomon@lineone.net geological survey of denmark and greenland bulletin 1, 61–73 (2003) © geus, 2003 62 tttt t t tttt t t t t t t t t t t t t t t t t t t t t t t t t t t t t t t t t t t t t t t t t t t t tt t t t t t t t t t t t t t t t t t t t t t t t t t t t t t t t t t t t t t t ttt t t t andøya 500 km vi ki ng s tra its m ar in e bo re al ba th on ian te thy s marine marine sm ar delt aic hrl ? ? ? ? nb lu aq ? vo he vi hp ? hi tr 211/21-1 ndb sp vg lbl co pb bo fr sw mc cnsd mo cg callo via n bbaajjoocciiaann -bbaatthhoonn iiaann b ajocian–bathonian carb onates jameson land east greenland ttt land paralic/deltaic siliciclastics volcanics coastline probable coastline speculative facies boundary tethyan shelf edge extent of the central north sea dome well 50° 40° 30° 0° 10°10° mnsh rfh skåne 9/10b-1 edge of that system unrivalled in detail. this is particularly the case in the middle jurassic. it provides the primary standard for comparisons with the rest of the world. to this denmark has contributed little, for its jurassic lies almost wholly in the subsurface. more recently, the post-war quest for petroleum has opened up what had largely been an immense geological terra incognita, the circum-global arctic. it quickly became apparent that the middle jurassic of the arctic could not be simply correlated with that of western europe because of non-overlapping endemisms in the distributions of the critical geochronometers, here the ammonites as guide-fossils. but at the same time, danish geological expeditions to east greenland revealed the key to correlations through the presence there of a residual sedimentary and biostratigraphical record so good that it has in turn become the chronostratigraphical standard for the whole of the arctic. this gives the jurassic of east greenland an importance out of all proportion to the modest extent of its outcrops. the middle jurassic makes also an important contribution to another area of more recent geological exploration. it is that of the former shelf-seas whose sediments now lie under the north sea and along the margins of the southern north atlantic, and whose interplay of tectonics and patterns of sedimentation form the major topics in other papers of the present volume. here, denmark does play a significant role by virtue of its position between the centres of active rifting in the north sea and the passive littoral margins of the baltic craton. the role of east greenland is twofold. firstly, it offers unique opportunities for direct study of fine examples, now splendidly exhumed onshore, of the kind of structures and basin-fills widespread in the subsurface of the north sea, important in the search for petroleum but necessarily mapped largely by indirect geophysical methods. secondly, the biostratigraphy of its ammonites provides the most precise available timecontrol on the genetic sedimentology and basin analysis, at a level of time-resolution that can significantly steer the interpretations. this chronostratigraphy is outlined here in some detail. palaeogeography, sedimentary basins and tectonics the middle jurassic experienced a continuing first-order worldwide (eustatic) rise in average sea level, reflected in the transgressive sediments still preserved on almost all of the surviving palaeocratonic, peri-oceanic margins of the former supercontinent pangea. taking the most optimistic estimate (see, for example, discussion by hallam 1988) of, say, + 200 m for the jurassic as a whole, this would suggest a mean value in the region of 50–100 m for thicknesses of epicontinental marine middle jurassic sediments that could be ascribed to eustatically-created accommodation space. one of the most plausible direct estimates to date (sahagian et al. 1996), based on the truly cratonic russian platform, gives about 30 m for the bajocian–callovian inclusive. whatever the meaning of such a figure might be generally, it does provide a yardstick when trying to estimate, or derive, the relative importance of the shorter-period, higher-order processes that modulate almost all sedimentary successions more locally. this includes the relative importance of higher-order eustatic fluctuations and regional tectonics in particular. palaeogeographically, the region around denmark during the jurassic lay still firmly inland on the laurentian craton (fig. 1). the landscape was one of gently undulating post-caledonian and post-hercynian topography whose highs had been eroded during the carboniferous and permian and whose lows had accumulated shallow marine evaporites and lacustrine or fluviatile continental sediments during the triassic, a period of historically low worldwide sea level. the marine transgressions marking the onset of a new major cycle of eustatic sea-level rise at the beginning of the jurassic then flooded a vast peneplain and inaugurated the regime of very extensive, shallow epicontinental seas so characteristic of the jurassic and cretaceous, far distant from the oceans and having few, if any, close current analogues. to the east lay the eurasian continent, almost as large as it is today. to the north lay fen63 facing page: fig. 1. the palaeogeography of western europe in the middle jurassic in early–mid bathonian times. the dot-dashed lines south of ar–lbl represent the approximate northerly limits of dominantly carbonate lithologies. the channel shown between greenland and the hatton–rockall landmass (the east greenland rift of ziegler 1982) is highly conjectural. aq, aquitaine basin; ar, armorica; bo, bohemia; cnsd, central north sea dome; co, cornubia; fr, franconia; he, helvetic highs (aar, gothard, mt. blanc massifs); hi, hurry inlet; hp, hebridean platform; hrl, hatton–rockall landmass; lbl, london– brabant landmass; lu, lusitanian basins; mc, massif central; mnsh, mid north sea high; mo, moray firth basin; nb, newfoundland banks; pb, paris basin; rfh, ringkøbing–fyn high; sm, spanish meseta; sp, shetland platform; sw, swabia; tr, trondheim fjord; vg, viking graben; vi, vindelicia; vo, vocontian trough. noscandia, firmly emergent during the whole of the jurassic. to the west lay the shallow seas covering the north sea, the british isles and what are now the continental shelves of the north atlantic, with the north american continent as far as idaho and british columbia beyond. to the south lay the shallow basins of the european platform, such the anglo-parisian, aquitanian, keltiberian, rhodano-helvetic, franco-swabian, lower saxonian and polish basins. the nearest ocean was the tethys, 1200 km to the south. access to it was not unrestricted but open through several broad channels. the only other access to the world’s oceans was northwards, through the viking straits, another broad seaway lying between fennoscandian norway and laurentian greenland (not the viking graben, a tectonic entity), connecting the central european seas with those of the arctic and thence the pacific. a number of residual hercynian rumps persisted as basin highs or even islands during the middle jurassic. they include the anglo-brabant landmass (‘london– ardennes island’), the scottish highlands, cornubia, armorica, the massif central, the harz (hercynia), bohemia and silesia. none of these were large or high enough to be major passive sources of primary coarse siliciclastic sediment, and although transgressive middle jurassic shoreline deposits may be observed at various places on them, the extent of their proximal facies is usually rather local. the largest source of both coarse and fine siliciclastics must have been the scandinavian crystalline massifs to the north and east and, in the viking straits, also greenland. the coarser sediments now occupy much of the shelves of the north atlantic and a broad belt extending from the northern north sea over denmark into the baltic. the finer sediments were very widely dispersed and the predominant facies of the european middle jurassic – the brown jura – north of the carbonate belts are clays, silts and fine-grained sands. supply in most cases kept up with demand and even in troughs created by local subsidence (see below), water depths stayed more or less constant. the epicontinental european seas were shallow, perhaps in the range 0–100 m, and quite small oscillations of relative sea level could produce transgressions and regressions over great distances. the positions of former shorelines are therefore often more than usually elusive. this general first-order picture was however widely disrupted by tectonic movements. the break-up of pangea had commenced and by the middle jurassic, gondwana had parted from laurasia. the north atlantic remained closed, but the central atlantic had opened and west africa was now separated from appalachian america by 1000 km of ocean. the effects of these major continental movements were clearly not localised in the tethys but felt also further north. one widespread expression was in the form of rifting into horst and graben or halfgraben structures. thus, the vocontian trough in southern france contains up to 750 m of middle jurassic sediments (terres noires); the lusitanian basin off western portugal up to 700 m; the central graben of the north sea, up to 1000 m; the viking graben of the northern north sea, up to 250 m; and the basins of the viking straits whose fill is now preserved onshore in east greenland, up to 700 m. these were correlated major perturbations extending over 2000 km on a single plate. extensional rifting was complemented by other perturbations of a different character: a transient regional uplift centred in the central north sea (ziegler 1982), driven by thermal upwelling of a subcrustal plume and followed by its dissipation and subsidence. the history of this event has been recently described in considerable detail by underhill & partington (1993, 1994). the rise of the dome began in the late early jurassic (toarcian). it broke marine surface in the aalenian with the onset of erosion, providing a new source of sediment for its surrounding basins. it culminated with the effusion at its centre of a thick lava pile up to 1500 m thick, probably during the bathonian–callovian. formerly termed the rattray formation (deegan & scull 1977), these middle jurassic volcanics have been redefined as the rattray volcanics member and the ron volcanics member, both of the pentland formation (richards et al. 1993). deflationary subsidence with erosional peneplanation was substantially complete in the late oxfordian and was followed by marine transgression. the differential uplift of the central dome is estimated to have been (at least) 400–500 m. the whole cycle of topographic uplift and subsidence had a period of 40 ma, but its highest point lay in the middle jurassic. its palaeogeographic effects extended over an area more than 1200 km across, from the east shetland platform in the north to the scottish highlands and yorkshire in the west, to the london–brabant massif in the south and the ringkøbing–fyn high in the east and taking in the mid north sea high – covering in fact almost the whole of the north sea. movements were more likely to have been pulsed than steady, both of doming and of rifting. their sedimentary expressions could therefore well have been at least partly time-correlated across the region and given rise to widely recognisable successions of sequences and systems tracts. but with such a strong tectonic overprint, both of doming and of rifting, claims to be able successfully to identify an 64 underlying higher-order eustatic cyclicity in the middle jurassic of the north sea seem somewhat improbable. palaeolatitude, climatology and bioprovincialism the shelf-seas of the european middle jurassic ranged in latitudes from about palaeolatitude 28°n (henceforth abbreviated to p-48°) in the south, at the northern margins of the tethys, to about p-48°n at, say, central east greenland (smith et al. 1981). the climatic gradient is clearly evident in sedimentary facies. in the south, carbonates dominate, such as lime mudstones and skeletal grainstones, including coarse echinoderm ‘breccias’. there are widely developed carbonate platforms and ramps. in the north, carbonates are rare or absent and never rock-forming. the pattern between these extremes is complex. it was evidently determined by an interplay of current patterns and topography, which strongly influenced the distribution of both benthic and nektoplanktonic biotas. thus, for instance, the eastern paris basin, the jura and what are now the helvetic alps (p-32°n) saw the build-up of a range of thick bioclastic carbonate platforms in the bajocian–bathonian. these shielded the swabo-franconian basin lying behind them from tethyan influence on its western side, its southern and eastern sides being bounded by the vindelician and bohemian massifs. the contrast between the paris and swabo-franconian basins, lying at similar palaeolatitudes, could therefore hardly be greater, the former abounding in a rich and highly diverse warmwater fossil biota; the latter, a sediment-starved basin with a highly impoverished fauna and almost no carbonates. further north-west, the warm-water carbonates and their rich biotas, with corals, reached the cotswolds of england (p-37°n), whereas the middle jurassic of northern germany at the same latitude was almost as impoverished as that of franconia. the furthest north that tethyan warm-water influences are detected is scoresby sund, east greenland (p-48°n), where bathonian hermatypic corals have been found very close to the transgressive western shoreline of the viking straits (håkansson et al. 1971; callomon & birkelund 1980). of immediate concern is the influence of the factors outlined above on the distribution of one group of organisms, the ammonites, for these are the leading guide-fossils whose biostratigraphy provides the primary standard chronostratigraphy over the whole region. the middle jurassic saw a strong segregation of the ammonites into three faunal provinces. these occupy successive, broadly latitudinal belts. the most southerly belt lies along the northern margins of the tethys, including iberia (but not the betic and subbetic basins of andalucia), aquitaine, the southern paris basin, the southern jura, peri-carpathian poland and much of the balkans. it is usually referred to as the submediterranean province of the tethyan realm. its ammonite faunas are rich, diverse and characteristic of warm-water, shallow neritic shelf-seas. they differ from those of the tethyan province proper mainly in compositions, which are reflections of different bioecological habitats. the true tethyan faunas are made up predominantly of pelagic groups such as phylloceras and lytoceras and are found in the pelagic carbonates of former seamounts and carbonate platforms still preserved, for example, in andalucia, sicily, the apennines, the venetian and austrian alps and the southern balkans. their places of entombment were not necessarily in deep water – some of them accumulated cross-bedded oolites – although their living habitats almost certainly were. they do occasionally mingle with the faunas of the shelf-seas, but then invariably as rare stragglers. the next belt constitutes the subboreal province of the boreal realm. it extended from the scottish highlands in the north (p-42°n) across northern europe (southern england, normandy, boulonnais, germany and the northern jura, northern and central poland) into the moscow–volga–donets basins on the russian platform, across the northern caucasus into trans-caspian turkmenistan and the basins of the amudarya, north of afghanistan, before now being lost under the himalayan ranges of the pamirs. in the aalenian, bajocian and bathonian, the distinction between submediterranean and subboreal faunas lies almost wholly in their diversities rather than in their mutually exclusive distributions. the provinces are therefore rarely explicitly differentiated in these stages and their faunas usually referred to simply as ‘european’. the diversity-differentiation is however already clearly discernible in the aalenian, during which typically submediterranean families reached dorset in abundance (callomon & chandler 1994) while being virtually unknown in swabia and franconia. the subboreal province became clearly differentiated from the submediterranean province only in the later middle jurassic, the callovian. it was characterised mainly by two families, the cardioceratidae and kosmoceratidae, whose distributions were complementary to those of the submediterranean reineckeiidae and oppeliidae. the boundaries could be sharp, as in going from southern normandy into touraine and poitou, or very diffuse, with much overlap, as in germany. 65 the third, most northerly, belt constitutes the boreal province of the boreal realm. it existed sporadically already in the toarcian and aalenian, but became clearly and strongly differentiated in the late early and early late bajocian. it then persisted into the cretaceous. it occupied the shelf-seas of the arctic (there being no positive evidence of the existence at that time of a truly oceanic arctic ocean), from the alaskan north slope through the yukon over the canadian arctic archipelago into svalbard and the whole of the barents sea including franz josef land, novaya zemlya and the petshora basin, the laptyev sea and kheta and lena basins in north-central siberia, and thence into north-eastern siberia and back to the bering straits. during the middle jurassic this sea, forming a rough triangle of sides 5000 km, was largely land-locked on two sides, by laurentia and eurasia (see map in callomon 1985, text-fig. 6a). communication in the early middle jurassic with the rest of the world’s oceans was possible in only two areas. the first was through a series of gateways into the northern pacific, between kolyma in far-eastern siberia and eastern alaska (although an anywhere near precise knowledge of where these gateways lay must await a satisfactory plate-tectonic reconstruction of this whole region). the second was the relatively narrow channel of the viking straits leading southwards to the epicontinental seas of europe and thence to the tethys. a third route opened in the late middle jurassic, in the callovian, when marine transgression from the petshora across the russian platform finally inundated the moscow–volga basins and thence completed the connections with the european epeiric seas from the east. the connecting route of particular interest here is the one between laurentia and fennoscandia along the viking straits. the largely landlocked middle jurassic seas of the arctic developed their own endemic faunas of ammonites, which lived and evolved there in almost total isolation from those further south. they did however venture down the viking straits as far as the shetland islands. here the seas divided into two channels (fig. 1). the westerly one passed southwards to the west of ireland, along the faroe–rockall troughs and perhaps others even further west, and gave connections with the western tethys presumably via the lusitanian seas west of iberia. little detail is so far known about the middle jurassic history of these channels. the easterly channel passed to the east of the shetlands in the general direction of the north sea and its sediments are preserved and well-documented in the viking graben, particularly in the brent field (deegan & scull 1977; underhill & partington 1993). the most southerly point from which fully marine boreal bathonian sediments with ammonites have so far been described appears to be at total’s well 9/10b-1 (bruce field, 59°40´n, p44°n; callomon 1979), 170 km ese of lerwick in the shetlands. progress further southwards into the north sea was then blocked by the emergent central north sea dome. its western non-marine flanking sediments have long been recognised in the various ‘estuarine series’ found in the aalenian–bathonian middle jurassic of britain, from scotland to the eastern midlands. marine connections between the northern north sea and the normal marine seas of wessex, and the north german and polish basins were completely cut by a physical barrier. no middle jurassic ammonites have been recorded from denmark, although typical european faunas are again well-represented in the southern baltic (stoll 1934). the ammonite faunas of the boreal and european provinces were wholly segregated: no region of overlap is known. both faunas evolved independently and provide the biochronologies on which the standard chronozonations are based. this means, however, that these chronozonations have had to be worked out independently, one standard scale for each province (see below), and correlations between these provincial scales continue to range from problematic to impossible. of particular interest here is the fact that the most detailed biochronology of the boreal bajocian–bathonian ammonites anywhere in the arctic has been that recorded in central east greenland, through an unusually favourable combination of circumstances (callomon 1993). it provides currently the chronostratigraphic standard of reference for the whole of the arctic. what of the channels to the west of ireland? no boreal middle jurassic ammonites have so far been found anywhere along their former courses, although this may be due wholly to lack of outcrop or sampling in boreholes. certainly no ammonites penetrated as far south as the lusitanian basins (p-30°n). yet that there must have been a fully marine connection is attested by those hermatypic corals found in milne land, east greenland: they could have got there by no other route. the failure of the ammonites to migrate southwards therefore probably reflects an ecological barrier: a case perhaps of inability to migrate against a northerly-flowing warm-water current carrying an inimical food-chain. direct marine connections between the arctic and europe were restored in the early callovian. a second boreal family of ammonites, the kosmoceratidae, also with a long pre-callovian history in the arctic, suddenly appeared in europe. but whether the connection was via the viking straits or via the russian platform 66 was until recently not clear: new discoveries now make the latter the more probable (gulyaev & kiselev 1999a, b; mitta 2000; mitta & starodubtseva 2000). the earliest, a kepplerites, has been found in the upper bathonian of franconia (dietl & callomon 1988). it resembles a form from the lower calyx zone of east greenland (see below), also of late bathonian age. the first horizon with abundant kepplerites keppleri (oppel) has been traced from southern england (upper cornbrash) via northern and southern germany to the northern jura and probably as far as the caucasus. it marks the base of the callovian stage and also has close relatives in east greenland. it is closely followed by the earliest horizon with abundant cadoceras of the cardioceratidae, also traceable from dorset via germany to the northern jura, with now closely related forms all over the russian platform (cadoceras elatmae (nikitin)). yet the earliest known ammonites from marine sediments directly overlying non-marine deltaic facies in scotland came from considerably higher levels, in the early koenigi zone on the east coast (moray firth) and in the late koenigi zone on the west coast (skye). furthermore, no callovian ammonites have been recorded from anywhere in the central north sea. indisputably direct marine connections between the arctic seas and those of europe via the viking straits were however fully reestablished in the late callovian (athleta zone). the same standard zonation can be applied to sediments from the top of the lower callovian upwards in both the viking straits as preserved in east greenland and throughout northern europe (see below). standard zonal chronostratigraphy the middle jurassic is made up chronostratigraphically of four standard stages: aalenian (lowermost, aalen, south-west germany), bajocian (bayeux, normandy), bathonian (bath, england) and callovian (uppermost, kellaways, wiltshire, england) (fig. 2). the chronometric ages of its boundaries are not directly very closely determined, but recent reviews (harland et al. 1990; odin 1994; gradstein et al. 1994) are in tolerable agreement in assigning to the middle jurassic a duration of 21 ma, i.e. close to a third of the 62–68 ma of the jurassic as a whole. published chronometric ages for individual stage boundaries are largely based on non-chronometric estimates of their relative durations, such as the number of standard chronostratigraphic zones or subzones they (now) contain, and should be treated with corresponding caution (see pálfy 1995; pálfy et al. 2000). the standard zones and subzones used here are chronozones defined by their bounding time-planes and are at the lowest two levels (vi and vii) of a hierarchy of successively finer subdivisions of the geological column, of the jurassic system (iii), its series (iv) and stages (v) (callomon 1994, 1995). the observational basis is the biostratigraphy of the most sensitively timediagnostic guide-fossils available, here the ammonites, which serve to characterise the zones in terms of their contents and to correlate successions. in most cases even finer time-resolutions are possible within a subzone, based on the recognition in it of more than one distinguishable – in an evolutionary sense – biohorizon, whose number can increase with the growth of knowledge of a highly incomplete geological record (callomon 1995). figure 3 shows standard zonations for the three ammonite biogeographic provinces represented in western europe. such zonations grow with time and knowledge. the oldest, going back substantially to oppel (1856–58) is that covering the northwest european or subboreal province, in which lie most of the outcrops studied during the classical age of geology (column a). it is the best known and most complete and hence provides the primary standard. the zonation used today in the submediterranean province is shown alongside it at the left (column b). its bathonian part became established in the 1940–50s, its callovian part a little later. the boreal zonation shown in the column at the right (c) was wholly unknown before 1959. the purpose of diagrams such as these is to represent what are effectively relative time-scales. but as there are no independent ways of estimating the relative durations of the time-periods represented by zones 67 oxfordian stagesseries callovian bathonian bajocian aalenian toarcian middle jurassic (dogger) fig. 2. the primary standard stages of the middle jurassic. 68 u pp er m id dl e lo w er u pp er m id dl e lo w er u pp er ba th on ia n c al lo vi an morrisi 37 36 35 33 32 31 34 30 29 28 24 23 22 21 20 19 18 17 14 13 11 10 9 8 3 2 1 ? ? zones zones & horizonszones lamberti athleta coronatum jason orbis zigzag parkinsoni garantiana subfurcatum zones calloviense koenigi herveyi discus nordenskjoeldi apertum calyx variabile cranocephaloide ishmae greenlandicus arcticus pompeckji indistinctus borealis lamberti (alligatus) athleta coronatum anceps gracilis herveyi (bullatum) discus retrocostatum zigzag bremeri subcontractus progracilis tenuiplicatus (as subboreal) (a) banksi polygyralis baculata dichotoma tetragona acris truellei bomfordi convergens macrescens yeovilensis tenuiplicatus progracilis subcontractus morrisi hodsoni blanazense hannoveranus hollandi discus keppleri terebratus kamptus gowerianus curtilobus galilaeii calloviense enodatum medea jason obductum grossouvrei phaeinum proniae spinosum henrici lamberti subzones & horizons subzones & horizons boreal province nw european / subboreal province submediterranean province b a c xx• xix• xviii• xvii• xvi• xv• xiv• xiiib• xiiia• xii• xi• xb• xa• ix• viii• viib• viia• vi• v• iv• iii• ii• i• 14• 13• 12• 11• 10• 9• ba jo ci an lo w er a al en ia n lo w er u pp er (submediterranean province not differentiated) concavum bradfordensis murchisonae discites ovalis sauzei humphriesianum laeviuscula formosum concavum gigantea bradfordensis murchisonae obtusum haugi sayni trigonalis laeviuscula cycloides humphriesianum blagdeni scissum opalinum fig. 3. the standard zonal chronostratigraphy of the middle jurassic in the three ammonite faunal biogeographic provinces represented in europe. the primary standard is the northwest european or subboreal, shown in the centre (column a). the secondary standard in the submediterranean province is on the left (b), that in the boreal province on the right (c). for annotations, see text. the submediterranean zonation for the aalenian and lower bajocian is the same as the northwest european/subboreal and is not separately indicated. no ammonite faunas of these ages are known from the boreal province itself. in the toarcian (lower jurassic) below, a separate zonation is again introduced for the mediterranean/tethyan province. or subzones, there is no unique or permanent way of drawing the diagrams. the scheme adopted here is to assume that the time-intervals between adjacent distinguishable ammonite biohorizons in the primary standard (a) are equal. as they are the smallest time-intervals at present resolvable by means of ammonites, the errors in this assumption are likely to have the least effect on the implied relative durations of the coarser units, the zones and subzones. the known number of such resolvable biohorizons is of course not final either. but as the european standard has been worked on intensively over a long time, it is safe to say that in the middle jurassic at least, we are now close to what is achievable in timeresolution by means of ammonites. taking the estimates of duration of the middle jurassic quoted at the beginning, of 21 (± 4?) ma, the average time-interval between faunal horizons (115) is 180 (± 36) ka; the average duration of a standard zone (28) is 750 (± 150) ka. correlation between european subboreal and submediterranean scales (a–b) is sufficiently close to draw them on the same scale. correlation between the european and boreal scales (a–c) even at zonal level is so imprecisely known that they have to be drawn quite independently. it is however probable that the lowest, borealis zone correlates with levels at around the lower–upper bajocian boundary. such a correlation had been previously postulated on the basis of a chain of correlations via the north american cordillera (callomon 1985, p. 64). it has now been confirmed through measurements of the stable isotopic ratio of strontium (87sr/86sr) in belemnites (j.h. callomon, unpublished data; m. engkilde, personal communication 1997). scale c is therefore also drawn on a faunal-horizon equal-interval assumption, based on the best available ammonite biochronology of east greenland. the boreal intervals have, however, been stretched somewhat relative to those of the european standard, to take the borealis zone down to the early upper bajocian. a: european subboreal primary standard the names of both zones and subzones as currently in use have been written out. to complete their definitions, the bounding time-planes of these standard chronostratigraphic units have to be fixed in boundary stratotype sections. while such definitions have been proposed for many of the zones and subzones, none have so far been formally ratified by the international commission on stratigraphy, whose aspirations do not currently descend below the level of standard stages (cowie et al. 1986). the lack of such formal definitions appears to have remarkably little consequence in practical day-to-day affairs. the reason has been pointed out repeatedly (callomon 1995) and is important. it is that rocks are correlated by means of the contents of zones – what lies between the bounding time-planes – rather than by correlation of the time-planes themselves. the wide variations in the relative time-values and numbers of subzones in zones as shown in the column are almost entirely of historical origin, reflecting a compromise between the growth of knowledge (faunal horizons) and a desire to retain some stability in zonal nomenclature. (in the times of oppel (1856–58) and of waagen (1867), the number of zones in what is today’s middle jurassic was 13; in arkell’s review of 1956 it was 21; here it has grown to 28). aalenian–bajocian based on recent revisions (callomon & chandler 1990; callomon 1995; callomon & cope 1995). bathonian the chronostratigraphy of this stage was one of the last to be worked out in europe because of exceptionally difficult ammonite biostratigraphy – scattered part-sections with only sparse ammonites. the first comprehensive modern syntheses by gabilly (1965) and mangold (1970) were based on the successions in western and eastern france, respectively. a review by westermann & callomon (1988) of its faunal successions was given in terms of 16 faunal horizons, some of which (9–14) are numbered here. some subsequent refinements in the upper bathonian of swabia were reported by dietl & callomon (1988), and a discrepancy between the middle and upper bathonian boundaries as drawn in the northwest european and submediterranean classifications was resolved by mangold (1988). the most recent reviews, by callomon & cope (1995, p.73) and mangold & rioult (1997), differ in some details of nomenclature but retain essentially the same biostratigraphic basis of faunal horizons. thus horizon 11, of procerites quercinus, is elevated into a quercinus subzone of the northwest european hodsoni zone, thereby raising the base of the orbis zone by one horizon; and the convergens and yeovilensis subzones of the zigzag zone are given separate names in the submediterranean province. the historically-conditioned difference between 69 the numbers of faunal horizons in the lower–middle and upper bathonian respectively is here particularly marked. callovian much new information in recent years has greatly amplified our knowledge of the lower and upper parts of the stage. the former was revised by callomon et al. (1988) and page (1989), the latter by dietl (1994 and unpublished data, based on excavations in swabia) and callomon & cope (1995). the herveyi zone is the former macrocephalus zone renamed, after it was discovered that the type-horizon of the index macrocephalites macrocephalus (schlotheim) did not in fact lie in its eponymous zone (callomon et al. 1992). a similar change of index will have to be made for the spinosum subzone of the athleta zone, because the type-horizon of kosmoceras spinosum (j. de c. sowerby) is now known to lie in the lamberti zone. only the coronatum zone remains in need of a modern revision. the middle–upper callovian boundary as reproduced here was defined in terms of the phyletic transients of the lineage of kosmoceras in england (callomon 1964, 1968) and not on the first appearance of peltoceras, which came in a little later. the upper boundary of the callovian stage, and hence the middle–upper jurassic boundary, has always been taken biostratigraphically to lie between the highest known faunal horizon of the lamberti zone, that of quenstedtoceras paucicostatum lange, and the lowest horizon of the oxfordian mariae zone, that of q. woodhamense arkell. it was usually easy to recognise because over large parts of northwest europe it is marked by a non-sequence. the discovery of expanded sequences across the boundary, e.g. in dorset and southern france, has made the final choice of stratotype section a matter of current debate. b: submediterranean secondary standard aalenian–bajocian the northwest european zonation is applicable and no separate scheme has been introduced. bathonian also reviewed by westerman & callomon (1988) and mangold & rioult (1997). the northwest european zonation is used for the lower and all but the topmost middle bathonian. differentiation became marked only in the upper bathonian. the submediterranean succession was worked out mainly in france, particularly in poitou (gabilly 1965), the southern paris basin and the southern jura (mangold 1970). there its uppermost part appears to be marked also by a widespread faunal or stratigraphical non-sequence, so that the bathonian– callovian boundary was until recently not clearly characterised (see below). callovian based predominantly on the faunal succession in western france worked out by cariou (1980, 1984). his labelling of faunal horizons (i–xx) has been included here, although the names of the subzones, many of which contain only a single faunal horizon, have been omitted. both bottom and top of the succession are marked by regionally widespread gaps. since figure 3 was drawn, however, a newly discovered section at buffevent, near niort in poitou, shows an expanded succession that spans the bathonian–callovian boundary (b. balusseau, p. branger and e. cariou, personal communications 2000). most particularly, one of its beds has yielded kepplerites keppleri, characteristic of the basal faunal horizon of the callovian stage by definition (callomon & dietl 2000). the bases of the callovian stages in both primary nw european (a) and secondary submediterranean (b) classifications can therefore be taken to coincide exactly. there remain some nomenclatural inconsistencies to be resolved, in that while some of the zones, e.g. gracilis and anceps, are named after distinct submediterranean indices, others, such as coronatum and athleta, share their names with their counterparts in the primary standard (a). yet these zones are independently defined and are of slightly different extents. the submediterranean athleta zone, for instance, has been chosen to begin at the first appearance of peltoceras in the hecticoceras trezeense horizon, xv, which lies at a level near the top of the phaeinum subzone of the primary athleta zone. the lowest callovian zone had for a long time been given a separate name and identity, the bullatum zone, as neither its base nor its top could be closely correlated with the primary standard. precise correlation of its base has now been achieved (see above) and so the main obstacle overcome. it has therefore been entered in figure 3 also under the name of the primary index as herveyi zone (olim macro70 cephalus). as its top, the base of the overlying gracilis zone, continues to be independently defined, it may be more appropriate, however, to continue to follow french usage (thierry et al. 1997) and to retain the bullatum zone in the submediterranean classification. an analogous alternative for the name of the top callovian zone in the submediterranean province, instead of lamberti, would be [perisphinctes] alligatus zone. c: boreal secondary standard this was recently reviewed by callomon (1993), based on the ammonite succession in east greenland. almost no ammonite faunas of aalenian – early bajocian ages have been recorded anywhere in the european arctic, and in much of the region the record of this period is marked either by stratigraphical gaps, or unfossiliferous or non-marine deposits (smelror 1994). in east greenland, the youngest ammonite faunas known from below the borealis zone are of early toarcian age (pseudolioceras lythense, bifrons zone). in svalbard and the barents shelf, the same is probably true, for claims to have recognized earliest aalenian ammonites, associated with toarcian forms, in a remanié conglomerate, the brentskardhaugen bed (wierzbowski et al. 1981), are based almost certainly on misidentifications as close aalenian homoeomorphs of the older toarcian species. the borealis zone is clearly recognisable throughout the whole of the arctic and is now known to correlate with the lower upper bajocian of europe or levels even a little older (see general discussion above). thereafter, the faunal horizons are numbered 1–35 up to the top of the lower callovian, as in callomon (1993), with a few subsequent additions based on recent unpublished field observations. they form the basis of the boreal zonation. zones have so far not been subdivided, and their different relative durations also reflect the historical evolution of knowledge. in the middle and upper callovian of east greenland, ammonites are so sparse (although the succession is thick) that they contribute nothing further to standard chronostratigraphy. elsewhere, for example in arctic russia and siberia, the upper callovian is divided into a lower zone of longaeviceras keyserlingi and an upper zone of eboraciceras subordinarium (meledina 1977), but the ammonite biostratigraphy is so fragmentary that they have not been included here. conclusion the middle jurassic of those parts of europe under discussion in this volume encompasses a highly diverse range of tectonic, sedimentary and palaeobiotic developments. to unravel their evolution successfully calls for the closest possible time-control of the rocks: of time-correlations over distances and of time-resolution more locally. fortunately, such time-control is provided by an abundance of good guide-fossils, the ammonites, whose biostratigraphy has given us a chronostratigraphic calendar of unrivalled finesse and hence, conversely, the clocks for dating individual rocks with great precision. the construction of the calendar has taken much effort over many years in many countries and the calendar that has emerged may seem complicated. but the rewards justify the effort needed to master these complexities. acknowledgement this review was written during the tenure of a leverhulme emeritus fellowship, which is gratefully acknowledged. references arkell, w.j. 1956: jurassic geology of the world, 806 pp. edinburgh, london: oliver & boyd. callomon, j.h. 1964: notes on the callovian and oxfordian stages. in: maubeuge, p.l. (ed.): colloque du jurassique à luxembourg 1962, 269–291. luxembourg: publication de l’institut grandducal. callomon, j.h. 1968: the kellaways beds and the oxford clay. in: sylvester-bradley, p.c. & ford, t.d. (eds): the geology of the east midlands, 264–290. leicester, uk: leicester university press. callomon, j.h. 1979: marine boreal bathonian fossils from the northern north sea and their palaeogeographical significance. proceedings of the geologists’ association (london) 90, 163–169. callomon, j.h. 1985: the evolution of the jurassic ammonite family cardioceratidae. in: cope, j.c.w. & skelton, p.w. (eds): evolutionary case histories from the fossil record. special papers in palaeontology 33, 49–90. london: palaeontological association. callomon, j.h. 1993: the ammonite succession in the middle jurassic of east greenland. bulletin of the geological society of denmark 40, 83–113. callomon, j.h. 1994: jurassic ammonite biochronology of greenland and the arctic. bulletin of the geological society of denmark 41, 128–137. 71 72 callomon, j.h. 1995: time from fossils: s.s. buckman and jurassic high-resolution geochronology. in: le bas, m. (ed.): milestones in geology. geological society memoir (london) 16, 127–150. callomon, j.h. & birkelund, t. 1980: the jurassic transgression and the mid–late jurassic succession in milne land, central east greenland. geological magazine 117, 211–226. callomon, j.h. & chandler, r.b. 1990: a review of the ammonite horizons of the aalenian – lower bajocian stages in the middle jurassic of southern england. in: cresta, s. & pavia, g. (eds): atti del meeting sulla stratigrafia del baiociano. memorie descrittive della carta geologica d’italia 40, 85–111. callomon, j.h. & chandler, r.b. 1994: some early middle jurassic ammonites of tethyan affinities from the aalenian of southern england. in: pallini, g. (ed.): proceedings of the 3rd pergola international symposium ‘fossili, evoluzione, ambiente’. palaeopelagos special publication 1, 17–40. rome: università ‘la sapienza’. callomon, j.h. & cope, j.c.w. 1995: the jurassic geology of dorset. in: taylor, p.d. (ed.): field geology of the british jurassic, 51–103. london: geological society. callomon, j.h. & dietl, g. 2000: on the proposed basal boundary stratotype (gssp) of the middle jurassic callovian stage. in: hall, r.l. & smith, p.l. 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(eds): siliciclastic sequence stratigraphy: recent developments and applications. american association of petroleum geologists memoir 58, 449–484. waagen, w. 1867: über die zone des ammonites sowerbyi. benecke’s geognostisch-paläontologische beiträge 1, 507–668. westermann, g.e.g. & callomon, j.h. 1988: the macrocephalitinae and associated bathonian and early callovian (jurassic) ammonoids of the sula islands and new guinea. palaeontographica a 203, 1–90. wierzbowski, a., kulicki, c. & pugaczewska, h. 1981: fauna and stratigraphy of the uppermost triassic and the toarcian and aalenian deposits in the sassenfjorden, spitsbergen. acta palaeontologica polonica 26, 195–241. ziegler, p.a. 1982: geological atlas of western and central europe, 130 pp. the hague: elsevier for shell internationale petroleum maatschappij. 73 manuscript received 19 april 1996; revision accepted 13 november 1998. research article alsen et al. 2020: geus bulletin 44. 5342. https://doi.org/10.34194/geusb.v44.5342 1 of 18 middle jurassic sandstone deposition in the wandel sea basin: evidence from cardioceratid and kosmoceratid ammonites in the mågensfjeld formation in kilen, north greenland peter alsen*1 , jussi hovikoski1 , kristian svennevig1 1geological survey of denmark and greenland (geus), copenhagen, denmark abstract age assessments from both palynostratigraphy and macrofossil biostratigraphy of the sandstone-dominated mågensfjeld formation, wandel sea basin, north greenland were hitherto hampered by post-burial thermal degradation of dinoflagellate cysts and a lack of well-preserved macrofossils. the formation was previously assigned to the upper cretaceous based on erroneous fossil identifications. finds of cardioceratid and kosmoceratid ammonites during recent field work now provide the first age control of the unit, demonstrating it to be of late bajocian – late bathonian and perhaps callovian (middle jurassic) age. this makes it among the oldest jurassic units, perhaps even mesozoic units, recorded in kilen, north greenland and eastern north greenland. previously, the complex structural and tectonic evolution of the area was poorly understood, and the structural relation of the mågensfjeld formation to the surrounding mesozoic units was a puzzle. the new age assessment simplifies the structural situation in the area significantly. further, the inference of a large reverse fault previously required to explain the proximity of the mågensfjeld formation to neighbouring jurassic units is now unnecessary. the data show that the wandel sea basin was influenced by the middle jurassic transgression and had sufficient accommodation space for marine deposition earlier than previously thought. the unit serves as a key datapoint and analogue for possible middle jurassic units in adjacent offshore basins. 1 introduction the first ever fossils reported from kilen in eastern north greenland (fig. 1) were collected by the greenarctic consortium and included the ammonite cranocephalites vulgaris, spath and various bivalves (dawes 1976; dawes & peel 1981). the fossils indicated the presence of a middle jurassic succession, but their exact location was not known. at this time, the geology of kilen had been interpreted solely from aerial photos as proterozoic–palaeozoic strata deformed during the caledonian orogeny (haller 1970). hence, the discovery of mesozoic fossils prompted the need for further field observations in this remote and inaccessible area. during pioneering expeditions in 1980, 1985 and 1998, e. håkansson and co-workers collected lithological, *correspondence: pal@geus.dk received: 26 june 2020 accepted: 17 sept 2020 published: 21 dec 2020 keywords: ammonite stratigraphy, middle jurassic, north greenland, sandstone deposition, wandel sea basin abbreviations: nhmd: natural history museum of denmark mguh: museum geologica universitas hafniensis/geological museum type collection (copenhagen, denmark) fa: facies association geus bulletin is an open access, peerreviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: sofie lindström (geus, denmark) reviewed by: mikhail rogov (geological institute of ras, russia) and one anonymous reviewer. funding: see page 15 author contributions: see page 16 competing interests: see page 16 additional files: none provided https://doi.org/10.34194/geusb.v44.5342 https://orcid.org/0000-0001-6218-9054 https://orcid.org/0000-0001-6330-8713 https://orcid.org/0000-0003-3863-8096 mailto:pal@geus.dk alsen et al. 2020: geus bulletin 44. 5342. https://doi.org/10.34194/geusb.v44.5342 2 of 18 www.geusbul let in.org sedimentological and biostratigraphic data in kilen. they described an essentially upper jurassic – upper cretaceous succession (håkansson et al. 1981b, 1991, 1993; pedersen 1991; dypvik et al. 2002), whereas the “middle jurassic” ammonites reported by dawes (1976) were identified as mid-cretaceous (albian) anahoplites cf. daviesi ornata, spath and a gastroplitinid (birkelund & håkansson 1983). fig. 1 simplified geological map of kilen in the wandel sea basin. a: the wandel sea basin is a post-caledonian fault-bounded carboniferous–lower palaeogene sedimentary basin in eastern north greenland (dawes & soper 1973; håkansson & stemmerik 1989; stemmerik et al. 1998). it covers the geographical areas of holm land (hl), kronprins christian land (kcl) and eastern peary land. b: the location of the mågensfjeld formation outcrop locality and the area covered by fig. 2 are shown. modified from svennevig et al. (2016). https://doi.org/10.34194/geusb.v44.5342 http://www.geusbulletin.org alsen et al. 2020: geus bulletin 44. 5342. https://doi.org/10.34194/geusb.v44.5342 3 of 18 www.geusbul let in.org fig. 2 detailed geological map of northern kilen fjelde. the geological situation around the mågensfjeld formation outcrop and the position of figs 3 and 4 are shown. the latter indicates the location of the mågensfjeld type section, which was logged, and where ammonites were collected. the now redundant reverse fault suggested by pedersen (1991) and håkansson et al. (1993) is indicated by a red dashed line. modified from svennevig (2018). https://doi.org/10.34194/geusb.v44.5342 http://www.geusbulletin.org alsen et al. 2020: geus bulletin 44. 5342. https://doi.org/10.34194/geusb.v44.5342 4 of 18 www.geusbul let in.org during a recent field campaign in north greenland, new data were retrieved from kilen. this area is key for studying the jurassic–cretaceous of north greenland and its relation to the conjugate barents margin, both lithologically and structurally. in 2012 and 2013, field teams sampled and logged all the sedimentary units described by håkansson et al. (1981, 1991, 1993). the objective of the recent campaign was to refine the biostratigraphy and establish a formal lithostratigraphic framework for an improved understanding of the depositional evolution of the succession (bojesen-koefoed et al. 2014; svennevig et al. 2016; alsen et al. 2017; hovikoski et al. 2018). several cretaceous intervals and units were poorly dated or not dated at all. the succession has been subject to significant thermal influence, which has removed organic palynomorphs hampering the use of dinoflagellate cysts stratigraphy (håkansson & pedersen 2001; svennevig et al. 2017; pedersen et al. 2018). macrofossils are absent or scarce and poorly preserved in several units. only a few intervals and units have published biostratigraphic ages (birkelund & håkansson 1983), whereas more detailed biostratigraphy were presented in unpublished reports (håkansson 1994; håkansson et al. 1994). parallel to the biostratigraphic and lithostratigraphic work, the structural geology of kilen was revised based on mapping from oblique photogrammetry and field work (svennevig et al. 2015). a new structural model for kilen was published, in which it was shown that late cretaceous extension caused normal faulting followed by eurekan compression, which gave rise to folding and thrusting (svennevig et al. 2016, 2017, 2018). a thick sandstone succession in northwesternmost kilen fjelde (the northern part of kilen) is informally referred to as unit 7 by håkansson et al. (1993) and forms the best exposed sediment succession in kilen. the steep, vertical outcrops in cliffs with nesting ivory gulls gave the unit its name: mågensfjeld formation (fig. 2). in danish, mågensfjeld refers to “mountain of the seagull.” pioneer geologists working in the region used the name “fuglefjeld” (danish for bird cliff). the age of the unit was considered coniacian (late cretaceous) due to the identification of inoceramid bivalves at the base of the section, and common imprints after belemnites and allegedly late cretaceous baculitid ammonites (håkansson et al. 1994). a late cretaceous age of the mågensfjeld formation significantly added to the structural complexity of kilen fjelde, and a large reverse fault was introduced to explain its proximity to upper jurassic units (pedersen 1991; håkansson et al. 1993). in this paper, we report new finds of fossils from mågensfjeld in kilen, collected in 2013, which demonstrate a middle jurassic age of the mågensfjeld formation. the dating of the unit has significant importance for the understanding of the structural geology of the area. 2 geological setting kilen forms part of the wandel sea basin in north greenland. it is a fault-bound basin containing a carboniferous – lower palaeogene sedimentary succession (dawes & soper 1973; håkansson & stemmerik 1989; stemmerik et al. 1998). the succession in kilen is triassic – upper cretaceous (håkansson et al. 1993; alsen et al. 2017; hovikoski et al. 2018). the area was deformed by several structural events, namely latest cretaceous extensional faulting followed by compression of presumably paleocene age. it has been divided into two major thrust sheets, which contain numerous smaller fault-bound structural units (svennevig et al. 2016, 2018). as a result, strata are folded and lithostratigraphic units are commonly fault bound (svennevig et al. 2016; hovikoski et al. 2018). 3 locality the mågensfjeld formation is exposed in the innermost, north-western part of the ice-bounded semi-nunatak kilen, surrounded by the ice cap flade isblink. the type section is situated at the steep north-east-slope of mågensfjeld in the uppermost reaches of the hondal valley (figs 1–4; hovikoski et al. 2018; svennevig 2018). 4 materials and methods 4.1 3d mapping of mågensfjeld the structural mapping of the mågensfjeld area builds on the geological map of kilen (svennevig 2018), which was mapped mainly by oblique photogrammetry (svennevig et al. 2015). oblique photographs were taken by a handheld digital camera from a helicopter, and triangulated and georeferenced, so that visible geological features such as bedding and faults could be mapped as 3d polylines. the polylines were used to calculate strike and dip of bedding and faults. these parameters were imported into a 3d modelling software along with the 3d polylines, an unpublished digitised field map (pedersen 1991), a digital mågensfjeld formation fig. 3 view from helicopter of the mågensfjeld formation at the “bird cliff,” mågensfjeld, looking south. https://doi.org/10.34194/geusb.v44.5342 http://www.geusbulletin.org alsen et al. 2020: geus bulletin 44. 5342. https://doi.org/10.34194/geusb.v44.5342 5 of 18 www.geusbul let in.org elevation model and an orthophoto and georeferenced field observations. in the 3d modelling software, robust geological models (3d maps) were built and used for structural validation and to produce 2d maps and cross sections. 4.2 ammonites we were granted access to the ammonite collection established by the pioneer geologists working in the area. one sample from “fuglefjeld” (now mågensfjeld) was labelled “dissolved baculites in sandstone” (the original label is in danish), which is probably the reason for assigning the sample and the sandstone unit to the upper cretaceous (coniacian). however, detailed inspection of the deeper part of the imprint of the dissolved supposed baculitid fossil shows the presence of a cone-shaped mould, which is clearly the filling of a belemnite alveolus. it adds to our own field observations that belemnites are common in the mågensfjeld formation, particularly in the lower part of the section. the report of late cretaceous inoceramids is also erroneous, as the inoceramid bivalves, which are relatively common, belong to the genus retroceramus. hence, there are no indications of upper cretaceous. our own collection contains 35 ammonite specimens collected in 2013 from the type section of the mågensfjeld formation. some samples were found loose in the scree below the outcrop; other specimens were found in situ at various levels throughout the measured section (figs 4 and 5). most specimens are poorly preserved, broken and fragmented. the sandstone in which they are contained has undergone significant alteration through diagenesis with extensive quarts cementation. fossils are commonly partly dissolved. the fossil material, however, contains readily identifiable specimens, which we have compared to the jurassic ammonite reference collection curated by j.h. callomon between 1960 and 2010 (callomon 1961, 1985, 1993; callomon et al. 2015). the reference collection is housed at the geological museum, copenhagen, now part of the natural history museum of denmark (nhmd). specimens represent middle jurassic ammonites of the cardioceratid genera cranocephalites, arctocephalites, arcticoceras, cadoceras and of the kosmoceratid genus kepplerites. selected ammonite specimens from mågensfjeld formation in kilen are shown in figs 6 and 7. 5 palaeontology our field work resulted in the discovery of numerous ammonites in situ in a number of levels throughout the 0 m 15–10 m 32 m 45 m 72 m ~80 m fig. 4 view towards the west, showing the mågensfjeld formation type section at mågensfjeld. this is perpendicular to the view shown in fig. 3. dashed line indicates the logged section (jhov-5/2013). ammonite-yielding levels are indicated. https://doi.org/10.34194/geusb.v44.5342 http://www.geusbulletin.org alsen et al. 2020: geus bulletin 44. 5342. https://doi.org/10.34194/geusb.v44.5342 6 of 18 www.geusbul let in.org cl 0 2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 32 34 36 38 40 42 44 46 54 56 58 60 80 82 84 88 90 92 94 96 98 100 kepplerites cf. traillensis or chisikensis, (~j-27) cadoceras apertum (j-24–26) cadoceras calyx (j-23) arcticoceras [crassiplicatum]? (j-17) arctocephalites cf. arcticus, (j-9) arctocephalites ishmae, (j-16) cranocephalites cf. pompeckji or furcatus (~po-8 or po-9) 102 m c vcs vf f m 3 fa protected shoreface protected shoreface protected shoreface delta front delta front truncated top fluvial?, fluvio-tidal 3 3 3 2 2 1 poorly exposed poorly exposed ? ? lithology sedimentary structures fossils trace fossils sandstone cross-bedding planar lamination/bedding asymmetric ripple cross-stratification wave-ripple cross-lamination bioturbated fragment belemnite logs bivalve parallel lamination plant fragments ammonite chondrites diplocraterion habichi helminthopsis nereites palaeophycus tubularis burrow mottling phoebichnus rhizocorallium isp. skolithos isp. thalassinoides isp. siphonichnus trend of the coarsests grain size fraction fig. 5 measured sedimentological log of the type section of the mågensfjeld formation (jhov-5/2013). fa: facies association; po-8, po-9, j-9 to j-27: faunal horizons for north-east greenland (see fig. 8); po: pompeckji; j: jameson land. modified from hovikoski et al. (2018, fig. 8b). https://doi.org/10.34194/geusb.v44.5342 http://www.geusbulletin.org alsen et al. 2020: geus bulletin 44. 5342. https://doi.org/10.34194/geusb.v44.5342 7 of 18 www.geusbul let in.org c. 100-m thick measured section at mågensfjeld (fig. 5; chosen as the type section, see hovikoski et al. 2018, fig. 8b). figured specimens are assigned official numbers from the geological museum type collection (mguh), copenhagen, denmark. these specimens are held in the nhmd repository. the remaining collections are housed at the geological survey of denmark and greenland (geus) and are assigned samples codes with the prefix “geus.” in the following section, the following abbreviations apply: s.s: sensu stricto; ms: manuscriptum; spp. indet: species plurimae indeterminata; sp. aff.: species affinis; cf.: confer. 5.1 systematic taxonomy family cardioceratidae siemiradzki 1891 genus cranocephalites spath 1932 cranocephalites cf. pompeckji (madsen 1902) or c. furcatus spath 1932 figs 6a–f d f c a h b e i jjg g i 1 cm fig. 6 selected cardioceratid ammonites from the mågensfjeld formation. a–f: cranocephalites sp. cf. pompeckji or furcatus mguh 33462–33467 (all from geus 545584). g, g: side and ventral views, respectively, of arctocephalites sp. delicatus or arcticus mguh 33468 (from geus 545586). h: arcticoceras ishmae, mguh 33469 (from geus 545569). i, i and j, j: side and ventral views of two specimens of arcticoceras spp. mguh 33470 and 33471, respectively (from geus 545568). https://doi.org/10.34194/geusb.v44.5342 http://www.geusbulletin.org alsen et al. 2020: geus bulletin 44. 5342. https://doi.org/10.34194/geusb.v44.5342 8 of 18 www.geusbul let in.org material and horizon. 15 specimens (11 body fossil fragments, 4 imprints) illustrated specimens include mguh 33462–33467 (from geus 545584) found laterally at c. 10 m height in the measured section (jhov-5/2013; figs 4 and 5). description. assemblage of relatively involute to moderately involute forms. ribbing fairly dense with strong primaries that divide into two or three secondaries at approximately mid flank. occasional intercalatories. in late stages, ribbing becomes coarser and blunt and fades on the venter. larger specimens tend to develop an inflated body chamber, typical of cranocephalites. the reference to the group of c. pompeckji and furcatus rests on the tendency towards relatively flattened sides, and general resemblance of the size, ribbing and coiling compared with specimens in the reference collection. occurrence and stratigraphy. the group of c. pompeckji and furcatus is common fossils in central east and north-east greenland, in particular on jameson land (callomon 1993; callomon et al. 2015) and on traill ø. the group represents the lower faunal horizons po-8 and po-9 of the c. pompeckji zone (fig. 8; donovan 1953, 1957; alsen 2000; vosgerau et al. 2004). c. pompeckji and furcatus are also reported in siberia (voronets 1962; meledina 1973), novaya zemlya, northern russia (cherkesov & burdykina 1979) and alaska (imlay 1962). remarks. specimens of c. pompeckji zone ammonites were found loose (ex situ) in kilen and on peary land (nygaard 2003). genus arctocephalites spath 1928 arctocephalites cf. arcticus (newton 1897) figs 6g and g material and horizon. one specimen, mguh 33468 (from geus 545586), found loose in scree in the lower part of the section jhov-5/2013 (fig. 4). description. a medium-sized specimen with rounded inflated whorls. it is crushed and the umbilicus is not visible, but it is clearly an involute form and the umbilicus must have been minute. it resembles mature forms 1 cm fig. 7 kepplerites cf. traillensis or chisikensis from mågensfjeld formation, mguh 33472 (from geus 545591). https://doi.org/10.34194/geusb.v44.5342 http://www.geusbulletin.org alsen et al. 2020: geus bulletin 44. 5342. https://doi.org/10.34194/geusb.v44.5342 9 of 18 www.geusbul let in.org of the oldest cranocephalites. however, upon comparison with the reference collection, it shows great resemblance with variants of arctocephalites arcticus, which occur stratigraphically higher than the cranocephalites of the c. pompeckji zone found below (geus 545584, see previous description of c. cf. pompeckji or furcatus). occurrence. the specimen indicates faunal horizon j-9 of the lowermost bathonian a. arcticus zone (callomon 1993; callomon et al. 2015; fig. 8). the species a. arcticus occurs in high latitudes in franz josef land, northernmost russia (newton & teall 1897; whitfield 1906), yukon, canada (poulton 1987), sverdrup basin, arctic canada (frebold 1964), siberia (meledina 1973) and pechora, russia (mitta 2009). a. arcticus is also found further south in the northern north sea (callomon 1975). genus arcticoceras spath 1924 arcticoceras ishmae (keyserling 1846) fig. 6h material and horizon. one specimen, mguh 33469 (from geus 545569), found in situ laterally of section jhov5/2013, approximately level with the 10–15 m interval in the measured section (figs 4 and 5). we obtained a mould of the relatively well-preserved left side of the specimen. fig. 8 middle jurassic ammonite zonal scheme from north-east greenland. j-1–37: succession of ammonite faunal horizons recorded on jameson land by callomon (1993). the faunal horizons of the c. borealis – c. pompeckji zones (j-1–8) were subsequently revised into the bo-1–po-23 faunal horizons shown in the expanded column to the right (callomon et al. 2015). j: jameson land; bo 1–3: faunal horizons in the borealis zone; in 1–8: faunal horizons in the indistinctus zone; po-1–23: faunal horizons in the pompeckji zone. https://doi.org/10.34194/geusb.v44.5342 http://www.geusbulletin.org alsen et al. 2020: geus bulletin 44. 5342. https://doi.org/10.34194/geusb.v44.5342 10 of 18 www.geusbul let in.org description. medium-sized with little umbilicus, dense strong regular ribbing curving gently forward from the umbilical seam to the venter. primaries bifurcate into secondaries at the mid-flank. occurrence. the taxon is the index species of the a. ishmae zone and indicates the faunal horizon j-16 in north-east greenland. a. ishmae is common from jameson land to the island of store koldewey, 500 km south of kilen (callomon 1993; callomon et al. 2015; fig. 8). a. ishmae is recognised widely throughout the boreal from yukon, canada (poulton 1987), northern russia (e.g. sokolov 1912; meledina 1987; mitta et al. 2015), novaya zemlya and northern siberia (meledina 1973), central russia (mitta et al. 2011) and the northern north sea (callomon 1975). remarks. callomon (1993) distinguished between two slightly different forms, transients α and β, of a. ishmae – the latter being a. ishmae s.s. together with the species a. harlandi rawson, which occurs in svalbard, northernmost norway (rawson 1982; ershova 1983), they represent faunal horizons j-14–16 in east greenland. based on material from a condensed section on the russian platform, those three taxa were suggested to be synonymous and to represent the intraspecific variation of one single variable species (kiselev 2020a, 2020b). this is not proven for the succession in greenland, and in this work, we follow callomon (1993). arcticoceras [crassiplicatum ms] callomon? material and horizon. imprint of one specimen at 32 m in section jhov-5/2013 (figs 4 and 5). the imprint could not be collected from the section, so a rubber cast (geus 545588) was made. description. the imprint represents only a fragment of the side view of c. 1/5 of a whorl from a rather large form. the fragment is densely and fairly strongly ribbed on a broadly rounded whorl side. after comparison with specimens in the reference collection, it is tentatively referred to a. [crassiplicatum ms], a species, introduced by callomon (1993), that has neither been described nor figured and type specimens have not been chosen (see remarks). occurrence. a. [crassiplicatum ms] occurs in the faunal horizon j-17 in the top of the a. ishmae zone, middle bathonian (callomon 1993; callomon et al. 2015; fig. 8). remarks. the taxon a. [crassiplicatum ms] is one of a series of undescribed and unfigured ms taxa introduced by callomon (1993). he intended to properly establish them with formal descriptions later, along with detailed accounts of their levels in the east greenland faunal succession. callomon succeeded with respect to formalising the cranocephalites taxa posthumously (callomon et al. 2015). several other ms-taxa unfortunately remain undescribed, but are meticulously curated in the nhmd reference collection. they have been widely integrated in the literature; for example, a. [delicatus ms] is a marker fossil of a key surface in the biostratigraphic framework for sequence stratigraphic analysis in engkilde & surlyk (2003). arcticoceras? spp. indet. figs 6i, i, j and j material and horizon. three specimens (all from geus 545568), ex situ, found in the scree slope (c. 25 m below section jhov-5/2013) below the steep outcrops of mågensfjeld formation (fig. 4). illustrated specimens include mguh 33470–33471. description. specimens with projected, somewhat coarser ribbing than a. ishmae but less coarse than a. [crassiplicatum] and less involute than a. ishmae. ribbing fades or disappears fully, and the venter becomes smooth. specimens have been crushed, resulting in a distorted shape, but they appear to have been relatively flat or discoidal compared with a. ishmae. tentatively referred to arcticoceras. genus cadoceras fischer 1882 cadoceras cf. calyx spath 1932 material and horizon. one specimen, geus 545589, found at 45 m in section jhov-5/2013 (figs 4 and 5). it is a fragmented, poorly preserved imprint. description. the fragment represents part of a relatively large shell with a relatively open umbilicus. ribbing developed from forward curving bullae. coarse to very coarse distant ribs on a rounded flank, which appear to be from a rounded, close to spherical whorl. ornamentation fades and the venter appears smooth. resembles the typical spherical forms of c. calyx. the preservation leaves some uncertainty, hence the cf. occurrence. index species of the c. calyx zone and occurring in faunal horizon j-23 in north-east greenland (callomon 1993; callomon et al. 2015; fig. 8). kopik & wierzbowski (1988) recorded c. cf. victor from assemblage with both c. calyx and c. apertum zones ammonites in an ironstone layer in the janusfjellet formation on svalbard, norway. cadoceras victor was subsequently considered synonymous with c. calyx by callomon (1993). c. calyx was reported from east siberia by knyazev et al. (2009) and the russia platform (mitta 2005; kiselev & rogov 2007). https://doi.org/10.34194/geusb.v44.5342 http://www.geusbulletin.org alsen et al. 2020: geus bulletin 44. 5342. https://doi.org/10.34194/geusb.v44.5342 11 of 18 www.geusbul let in.org cadoceras apertum callomon & birkelund 1985 material and horizon. one specimen, geus 545590, found at 72 m in a poorly exposed part of the section jhov-5/2013 (figs 4 and 5). fragment. description. the fragment is c. 1/5 of a whorl. only the upper part of one side of the whorl and the venter are preserved. it is totally smooth. the venter is relatively narrow and arched. it does not resemble the thickwhorled holotype of c. apertum, but the reference collection includes smooth macroconch variants with slender, narrowly rounded venter. occurrence. index of the c. apertum zone in east greenland, faunal horizon j-24–26 (callomon 1993; callomon et al. 2015; fig. 8). outside greenland, the taxon occurs in the janusfjellet formation, svalbard (kopik & wierzbowski 1988) and russia (mitta 2005). family kosmoceratidae haug 1887 genus kepplerites neumayr & uhlig 1892 kepplerites cf. traillensis donovan 1953 or k. chisikensis imlay 1975 fig. 7 material and horizon. one specimen, mguh 33472 (from geus 545591), consists of part of the internal mould/ steinkern only and imprints of the rest of the specimen, which are crushed and incomplete. found at c. 80 m height in section jhov-5/2013 (figs 4 and 5). description. the last whorl is clearly uncoiling, showing the specimen to be an adult. maximum diameter is 131.5 mm. ribbing is fine and dense. primary ribs are relatively prominent compared with the secondaries. it begins slightly rursiradiate to rectiradiate before crossing the gently sloping umbilical wall and shoulder with a gentle forward bend and becomes slightly prorsiradiate. the projection increases near the end or final peristome. the primary ribs divide low, about one-third, up the flank into forward leaning sheaves of four straight, evenly spaced and equally strong secondaries. comparisons. the specimen is close to both k. traillensis and k. chisikensis. the latter includes kepplerites tenuifasciculatus described from east greenland (callomon 2004), but k. tenuifasciculatus was recently considered a junior synonym of k. chisikensis imlay by mönnig & dietl (2017). k. traillensis is less densely ribbed than k. chisikensis; at last septum, k. traillensis has 31 primaries per whorl, whereas k. chisikensis has 40 (callomon 2004). the collected specimen is too poorly preserved to be identified as either of those two species. hence, it is referred to as k. cf. traillensis or k. chisikensis. the species kepplerites vardekloeftensis callomon has coarser secondaries and ornamentation that fades on the last body chamber. adult kepplerites svalbardensis sokolov & bodylevsky is generally smaller. the studied specimen resembles k. svalbardensis in coiling and ribbing density. however, k. svalbardensis is smaller, and its ribbing differs in becoming backwards, curving on the upper flank, after the initial forward curving on the lower flank and after crossing the umbilical shoulder. lastly, the specimen from mågensfjeld is also close to some variants of k. keppleri (oppel) with relatively delicate primaries (e.g. quenstedt 1887, plate 77, fig. 3; tintant 1963, plate 1; page 1989, figs 5.1a, b) but differs from those with relatively coarser primaries (e.g. tintant 1963, plate 2; dietl & gygi 1998, plate 1a). occurrence. k. chisikensis and k. traillensis indicate the faunal horizons j-24–27 in the c. apertum zone (callomon 1993, 2004; fig. 8). note that this zone was initially considered as callovian, based on strong resemblance of kepplerites from this zone with k. keppleri, the appearance of which defines the base of the callovian in europe (callomon 1993; mönnig & dietl 2017). mitta & alsen (2013) tentatively placed the c. apertum zone in the upper bathonian, based on late bathonian ornamentation recorded in material above k. “tenuifasciculatus.” considering that ornamentation in the genus developed progressively or regressively (callomon 2004), thick ribs do not unequivocally demonstrate a bathonian age. reference of the c. apertum zone to the bathonian was supported by mönnig & dietl (2017). kepplerites dietli schairer is considered a junior synonym of k. traillensis that correlates the greenland faunal horizon j-26 of callomon (1993) in the c. apertum zone to the lower part of the c. discus zone in germany. the k. aff. traillensis in faunal horizon j-28 is considered synonymous with k. radiatus sakharov & lominadze and thus correlates to the upper part of the c. discus zone (mönnig & dietl 2017). the possible presence of k. keppleri in the faunal horizon j-29 in the nordenskjoeldi zone would provide correlation to the base callovian k. keppleri fauna, defining the base of the callovian in europe (mönnig & dietl 2017; fig. 8). if the present specimen is actually k. keppleri, which cannot be determined from its preservation, it would be indicative of the fauna j-29–30 in the nordenskjoeldi zone and the base callovian. this fits with the studied section where it was found somewhat above c. apertum of the apertum zone (fig. 5). 5.2 biostratigraphic summary the biostratigraphy at mågensfjeld can be summarised as follows: 1. late bajocian cranocephalites cf. pompeckji or c. furcatus represents the po-8 or po-9 faunal horizons of the c. pompeckji zone (callomon et al. 2015). 2. bathonian arctocephalites cf. arcticus represents the j-9 faunal horizon of the a. arcticus zone. https://doi.org/10.34194/geusb.v44.5342 http://www.geusbulletin.org alsen et al. 2020: geus bulletin 44. 5342. https://doi.org/10.34194/geusb.v44.5342 12 of 18 www.geusbul let in.org 3. bathonian arcticoceras ishmae and arcticoceras [crassiplicatum] of the j-16 and j-17 horizons represent the a. ishmae zone (callomon 1993; callomon et al. 2015). 4. the upper bathonian c. calyx and c. apertum zone are represented by a relatively thick interval ranging from the middle of the section to near the top of the outcrop. the ammonites include c. calyx, indicating the j-23 horizon, c. apertum of the j-24– 26 horizons and kepplerites cf. traillensis or chisikensis of the j-27 horizon (callomon 1993; callomon et al. 2015; fig. 7). nygaard (2003) recorded one loose cranocephalites aff. borealis found loose in kilen in an “area covered by upper cretaceous sediment” (nygaard 2003, p. 6). it was fig. 9 examples of facies from the mågensfjeld formation. a: base of the succession showing the contact between facies association (fa) 1–2. sb: possible sequence boundary; fs: flooding surface. b: typical coated expression of tabular beds of fa 2. c: parallel-laminated sandstone that grades upwards into ripple cross-lamination, fa 2. d–f: typical trace fossils of mågensfjeld formation. d: diplocraterion isp. (di) and skolithos isp. (sk) burrowing into burrow-mottled fine-grained sandstone. e: bedding plane view to radiating burrow arms of phoebichnus isp. f: rhizocorallium isp. in fine-grained sandstone. https://doi.org/10.34194/geusb.v44.5342 http://www.geusbulletin.org alsen et al. 2020: geus bulletin 44. 5342. https://doi.org/10.34194/geusb.v44.5342 13 of 18 www.geusbul let in.org likely found beneath mågensfjeld. the ammonite faunas described in the present study occur in the steep well-exposed cliffs in the top of the mågensfjeld, but the scree-covered lower flanks of the hill leave room for older deposits – and a cranocephalites borealis bearing horizon. 6 depositional setting the mågenfjeld formation outcrop is characterised by a weathered rock surface with a pale coating that commonly hinders detailed sedimentological data collection. tentatively, the succession is divided into three facies associations (fa 1–3; figs 8 and 9). 6.1 facies association 1: fluvial or fluviotidal? 6.1.1 description fa 1 has a single occurrence at the base of the mågensfjeld formation succession. the deposits are poorly exposed and only a few metres can be studied in any detail. the unit is sharply overlain by fa 2 (see section 6.2; fig. 9a), but the lower boundary was not observed. the observable facies includes a few centimetre-thick, unbioturbated, mud-clast-bearing, fine-grained sandstone beds and lenticular bedding that is rich in plant debris. the mud-clasts are commonly concentrated at the base of the bed. the top of fa 1 is demarcated by an erosion-based extraformational pebble bed, a few centimetres thick. 6.1.2 interpretation due to the limited data, we do not attempt to make a detailed palaeoenvironmental interpretation of fa 1, and the following discussion is considered tentative. lack of bioturbation in a periodically, relatively low-energy setting (e.g. post-event bed boundaries) may suggest that the endobenthic colonisation was limited by the freshwater setting. the upper erosional boundary and abrupt change in grain size from fine-grained sandstone to pebbles suggest truncation and development of an erosional lag. such a surface may form, for instance, as a result of lateral changes in channel position in fluvial, estuarine or deltaic environments, or as a result of a progradational jump (even forced regression), and therefore could represent a sequence boundary. furthermore, considering that the overlying deposits (fa 2) are marine, the surface is also associated with deepening. in summary, considering the aforementioned observations and the stratigraphic position below fa 2, the deposits are broadly interpreted to represent a fluvial or fluviotidal environment. the top of fa 1 represents a flooding surface and possibly a sequence boundary. 6.2 facies association 2: delta front 6.2.1 description fa 2 is a common facies association type in the lower and middle parts of the mågensfjeld formation. it forms aggradational to slightly upwards coarsening successions up to 10 m thick, which comprise decimetre-scale, tabular fine-grained sandstone beds (figs 9b and c). the lower contact of fa 2 is sharp, and the limited exposure suggests a down-lapping pattern onto fa 3 (see section 6.3) or fa 1. the upper contact to fa 3 is either sharp or gradational. the rock surface is usually covered hindering systematic data collection, but where observable, the beds are structureless or show parallel lamination capped with ripple-cross lamination (fig. 9c). the beds show straight or trough-shaped bases. the beds appear to pinch-out towards the south-east. similarly, sporadic palaeocurrent measurements tentatively suggest south– south-east oriented flow (n = 3). bioturbation is commonly sporadic (bioturbation index 0–2, locally higher) and concentrated towards the top of the beds (data not shown). the suites are of low diversity and consist predominantly of vertical to inclined trace fossils such as diplocraterion isp., skolithos isp. and ?siphonichnus isp. (fig. 9d). sub-millimetre to millimetre-scale mud-filled branching burrows (?chondrites) and indistinct burrow mottling are locally present. plant fragments are common and include complete leaves. reworked belemnites are locally concentrated at the base of the beds. 6.2.2 interpretation the trace fossil assemblage and the occurrence of marine fossils such as belemnites point to a general shallow marine setting. tabular beds showing parallel lamination to ripple cross lamination may suggest an origin of a shallow marine gravity flow, whereas wave-generated structures are rare. putative downlapping, the pinch-out nature of the beds and the presence of deposits, rich in plant material supports a distal delta front setting. 6.3 facies association 3: protected shoreface 6.3.1 description fa 3 is a common facies association type in the middle and upper parts of the formation. it overlies fa 2 sharply or gradationally and is sharply overlain by the same association, where observable. the association forms aggradational or faintly upwards coarsening successions, comprised of bioturbated fine-grained sandstone. distal expressions consist of intensively bioturbated muddy, very fineto fine-grained sandstone. common trace fossils include indistinct mud-filled grazing structures (?helminthopsis isp. and nereites isp.), mud-filled branching burrows (?chondrites), palaeophycus tubularis and ?siphonichnus isp. proximal expressions consist of fine-grained sandstone burrowed with a low to moderate diversity assemblage, showing recurring phoebichnus isp., rhizocorallium isp., thalassinoides isp., diplocraterion isp., skolithos isp. and ?siphonichnus isp. in addition to aforementioned structures (figs 9e and f). wave ripples and dune-scale cross bedding are rarely observed. https://doi.org/10.34194/geusb.v44.5342 http://www.geusbulletin.org alsen et al. 2020: geus bulletin 44. 5342. https://doi.org/10.34194/geusb.v44.5342 14 of 18 www.geusbul let in.org 6.3.2 interpretation the trace fossil content generally points to a relatively low-energy, shallow marine environment periodically influenced by small-scale event sedimentation. the high bioturbation intensity and scarcity of preserved wave or storm-generated structures probably point to a confined setting, where oscillation currents are subdued. fa 3 is burrowed with elements of the cruziana ichnofacies, with increasing contribution of elements of the skolithos ichnofacies towards proximal expressions and is interpreted to represent protected lower to middle shoreface-like environments. 6.4 summary in summary, fa 1 is recorded lowest in the measured section between 0 and 2 m, whereas the remaining overlying section belongs to alternations of fa 2 and fa 3. overall, the succession reflects a backstepping to aggradational depositional evolution with a putative fluvial or fluviotidal base, overlain by deltaic and protected shoreface facies. 7 discussion 7.1 stratigraphic relations in kilen the mågensfjeld formation is only well-exposed at mågensfjeld, on the upper slopes of the mountain. below the exposures, the slope is covered by scree composed of blocks and slabs of thermally altered sandstone weathered out from the mågensfjeld formation. the scree covers the boundary to the underlying gletscherport formation, which has a restricted exposure near the foot of the slope towards the hondal valley (hovikoski et al. 2018; fig. 2). the orientation of the strata of the mågensfjeld formation is close to horizontal (figs 3 and 4), whereas the strata of the gletscherport formation are steeply dipping (up to 70°) towards the east. this is most likely due to fault drag as the gletcherport formation is exposed just west of the inferred position of a large normal fault in the hondal valley (normal fault ii of svennevig et al. 2016). the fault has downthrown the midto upper cretaceous succession at saddelfjeld to the east, relative to the middle jurassic succession at mågensfjeld to the west. an alternative explanation for the marked difference in orientation between the mågensfjeld and the gletscherport formations could be an angular unconformity between the two units, covered by the scree, indicating a marked structural event between the depositions of the two units. however, we consider the fault drag explanation to be the most plausible. the upper boundary of the mågensfjeld formation is not exposed. the succeeding unit in the kilen succession is the birkelund fjeld formation, which has yielded kimmeridgian ammonites (hovikoski et al. 2018). the lower part of the birkelund fjeld formation is not exposed in outcrop, and the age of its base is thus unknown. this leaves open the possibility of a depositional gap between the two formations, or for the presence of callovian–oxfordian strata in the interval between the bathonian of the mågensfjeld formation and the kimmeridgian, upper part, of the birkelund fjeld formation. 7.2 middle jurassic in the neighbouring high arctic 7.2.1 north-east greenland the nearest middle jurassic ammonite-bearing succession in greenland is the pelion formation on store koldewey island (c. 76°n, north-east greenland), where a rich ammonite fauna represents the bathonian–callovian (ravn 1911; piasecki et al. 2004). the pelion formation is a widespread deltaic to shallow marine sandstone unit with outcrops south of store koldewey on hochstetter forland, kuhn ø, wollaston forland, hold with hope, geographical society ø, traill ø and as far south as jameson land (see surlyk 2003; surlyk et al. in press). a report of c. aff. borealis in kilen indicated that middle jurassic deposition commenced in the c. borealis chron (nygaard 2003). this would be isochronous with the onset of the huge marine flooding and deposition of the sand-dominated, shallow marine and deltaic system of the pelion formation (callomon 1993; surlyk 2003). the pelion formation probably extends offshore into the danmarkshavn basin north of store koldewey. 7.2.2 peary land middle jurassic deposits may be present on peary land, making them the closest to kilen. håkansson & pedersen (2015) subdivided the sedimentary succession of the wandel sea basin into a series of separate, discrete basins, many of which they considered to be formed by strike-slip movements within the kronprins christian land strike-slip mobile belt. in one restricted area of sildredome, eastern peary land (fig. 1), lies the so-called sildredome basin – so-named by håkansson & pedersen (2015). the area is a few square kilometres in size and contains a fault complex. numerous middle jurassic c. pompeckji zone ammonites have been recorded in the area (nygaard 2003; håkansson & pedersen 2015) and found loose on a weathered surface mapped as the ladegårdsåen formation, peary land (kokfelt et al. 2013). more than 20 such specimens have been counted from this small area, which is strongly indicative of the middle jurassic (upper bajocian). however, outcrops are yet to be discovered. the oldest age of the ladegårdsåen formation has so far been dated middle oxfordian (håkansson et al. 1981; dypvik et al. 2002). 7.2.3 svalbard today, svalbard and the barents shelf are more than 600 km away from kilen, but prior to the opening of the atlantic, they were only 400 and 200 km away, respectively https://doi.org/10.34194/geusb.v44.5342 http://www.geusbulletin.org alsen et al. 2020: geus bulletin 44. 5342. https://doi.org/10.34194/geusb.v44.5342 15 of 18 www.geusbul let in.org (svennevig et al. 2017, using müller et al. 2016). the closest (onshore) correlative unit to the wandel sea basin, geologically speaking, is probably found on svalbard. middle jurassic ammonites have been reported from svalbard in the oppdalen member of the lower part of the agardhfjellet formation, immediately above thin (decimetre to a few metres thick) remanié deposits of the brentskardhaugen bed (birkenmajer 1980; birkenmajer et al. 1982). the brentskardhaugen bed is a conglomerate with reworked concretions containing toarcian, aalenian and bajocian fossils (bäckström & nagy 1985). dypvik et al. (1991a, 1991b) considered the brentskardhaugen bed deposited close to the bathonian–callovian transition. however, that age is in conflict with the occurrence of cranocephaloide zone ammonites in the overlying agardhfjellet formation (kopik & wierzbowski 1988). the cranocephaloide zone is now considered middle bathonian (callomon et al. 2015). the agardhfjellet formation is up to 290 m thick and is bathonian to ryazanian. the lower oppdalen member is a silt and fine-sand-dominated unit, 10–60 m thick and relatively coarse compared to the dominantly papershales of the agardhfjellet formation (mørk et al. 1999). kopik & wierzbowski (1988) reported a succession of three assemblages from the lower 10 m of the oppdalen member of middle bathonian cranocephaloide zone, calyx–apertum zones and callovian coronatum–athleta zones, respectively, with several taxa related to those reported here from kilen, north greenland. the documentation of a middle jurassic sandstone unit in kilen provides an important data point for sandstone deposition and a reservoir unit analogue in the region. kilen is part of the same depositional system as the northern part of the store koldewey platform bordering the danmarkshavn basin in the offshore area east and north of store koldewey island. the danmarkshavn basin was recently subject to petroleum exploration with five licence blocks issued. the mågensfjeld formation may be a more closely fitting analogue for sandstone units in the western barents sea (e.g. stø formation; olaussen et al. 1984; gjelberg et al. 1987; klausen et al. 2019) than the onshore analogues on svalbard. the presence of bajocian–bathonian shallow marine deposits in kilen shows that the wandel sea basin was influenced by the middle jurassic transgression and had sufficient accommodation space for marine deposition earlier than previously thought (cf. dypvik et al. 2002). the palaeoenvironmental characteristics of the middle jurassic mågensfjeld formation and upper jurassic (kimmeridgian) birkelund fjeld formation in kilen are strikingly similar. both are comprised mainly of protected shallow marine and deltaic environments and show generally an aggradational stacking pattern. however, a deepening trend from a fluvial, fluviotidal and deltaic environment to a protected shoreface occurs in the bajocian interval. 7.3 structural implications the new middle jurassic age for the mågensfjeld formation greatly simplifies the structural complexity of the area. the previous late cretaceous age called for the introduction of a large inferred reverse fault, bringing the late jurassic kuglelejet formation in contact with the previously dated late cretaceous mågensfjeld formation (pedersen 1991; håkansson et al. 1993; fig. 2). using the estimated thicknesses of the formations (hovikoski et al. 2018; svennevig et al. 2016), the vertical throw on the suggested fault would have been around 2000 m – by far the largest known from the area. the age of the mågensfjeld formation presented here, along with mapping from oblique photogrammetry and 3d modelling (svennevig et al. 2015, 2016), demonstrate a simple, gentle and upright anticline with the middle jurassic mågensfjeld at the core, truncated by a normal fault to the east and covered by ice to the west and north (svennevig 2018; figs 1 and 2). the normal fault is presumably of latest cretaceous age and predates the eurekan folding that likely occurred during the paleocene–eocene (svennevig et al. 2016). 8 conclusions the sandstone succession exposed in the steep cliffs of mågensfjeld, belonging to the mågensfjeld formation, records generally shallow marine depositional environments. the succession is fossil bearing, and the ammonites belonging to the genera cranocephalites, arctocephalites, arcticoceras, cadoceras and kepplerites show a great affinity to the middle jurassic faunal succession in north-east greenland. the ammonites allow the identification of the bajocian c. pompeckji zone, the bathonian a. arcticus, a. ishmae, c. calyx and c. apertum zones. the middle jurassic age of the unit allows for a much more straightforward model of the structural evolution of this area of kilen, compared with the previously proposed late cretaceous age. the structural evolution essentially records late cretaceous extension followed by palaeogene folding (probably eurekan, paleocene–eocene). the deposition of the mågensfjeld formation records middle jurassic transgression in the wandel sea basin and the development of accommodation space for marine deposition of a sand-dominated unit. the age of the unit is equivalent to the well-described reservoir unit analogues of the middle jurassic in north-east greenland (e.g. pelion formation, store koldewey). as such, the mågensfjeld formation serves as a key datapoint and analogue for possible middle jurassic units in offshore basins; for example, in the western barents sea, the wandel sea and the danmarkshavn basins. acknowledgements we thank jette halskov for drafting of figs 1, 2, 4, 5 and 8. we appreciate the useful comments of the two reviewers, m. rogov and one anonymous reviewer, which greatly improved the final manuscript. https://doi.org/10.34194/geusb.v44.5342 http://www.geusbulletin.org alsen et al. 2020: geus bulletin 44. 5342. https://doi.org/10.34194/geusb.v44.5342 16 of 18 www.geusbul let in.org additional information funding statement data were collected during the field work funded by the geological survey of denmark and greenland (geus) and undertaken within the collaboration project “petroleum geological studies, services and data in east and north-east greenland.” author contributions pa: conceptualisation. pa, jh and ks: formal analysis, investigation, writing – original draft, writing – review and editing. competing interests the authors declare no competing interests additional files none provided. references alsen, p. 2000: middle jurassic ammonite biostratigraphy in the traill ø region north-east greenland, 104 pp. unpublished master’s thesis, university of copenhagen. alsen, p. et al. 2017: the isrand formation: a middle triassic daonella-bearing, black shale unit in kilen, north greenland (with a note on the triassic in amdrup land). newsletters in stratigraphy 50, 31–46. https://dx.doi.org/10.1127/nos/2016/0341 bäckström, s.a. & nagy, j. 1985: depositional history and fauna of a jurassic phosphorite conglomerate (the brentskardhaugen bed) in spitsbergen. norsk polarinstitutt skrifter 183, 61. birkelund, t. & håkansson, e. 1983: the cretaceous of north greenland – a stratigraphic and biogeographical analysis. zitteliana 10, 7–25. birkenmajer, k. 1980: jurassic–lower cretaceous succession at agardhbukta, east spitsbergen. studia geologica polonica 66, 35–52. birkenmajer, k., pugaczewska, h. & wierzbowski, a. 1982: the janusfjellet formation (jurassic–lower cretaceous) at myklegardfjellet, east spitsbergen. palaeontologia polonica 43, 107–140. bojesen-koefoed, j., alsen, p. & christiansen, f.g. 2014: six years of petroleum geological activities in north-east greenland (2008–2013): projects and a view of the future. geological survey of denmark and greenland bulletin 31, 59–62. https://dx.doi.org/10.34194/geusb. v31.4661 callomon, j.h. 1961: the jurassic system in east greenland. in: raasch, g.o (ed.): geology of the arctic, 258–268. toronto: university of toronto press. https://dx.doi.org/10.3138/9781487584979-024 callomon, j.h. 1975: jurassic ammonites from the northern north sea. norsk geologisk tidsskrift 55, 373–386. callomon, j.h. 1985: the evolution of the jurassic ammonite family cardioceratidae. in: cope, j.c.w. & skelton, p.w. 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https://dx.doi.org/10.17850/njg97-1-02 svennevig, k., et al. 2018: descriptive text to the geological map of greenland, 1:100 000, kilen 81 ø.1 syd. geological survey of denmark and greenland map series 8, 29 pp. + map. https://dx.doi.org/10.34194/ geusm.v8.4526 tintant (h.), 1963: les kosmocératidés du callovien inférieur et moyen d’europe occidentale. essai de paléontologie quantitative. publications de l’université de dijon, t. 29, 1–500. https://gallica.bnf.fr/ ark:/12148/bpt6k4807624c/f11.item.texteimage.zoom voronets, n.s. 1962: stratigraphy and cephalopod molluscs of the jurassic and lower cretaceous deposits of the lena-anabar region. trudy nauchno-issledovetel’skogo instituta geologii arktiki 110, 237 pp. (in russian) vosgerau, h., et al. 2004: jurassic syn-rift sedimentation on a seawards-tilted fault block, traill ø, north-east greenland. geological survey of denmark and greenland bulletin 5, 9–18. https://dx.doi. org/10.34194/geusb.v5.4800 whitfield, r.p. 1906: notes on some jurassic fossils from franz josef land, brought back by a member of the ziegler exploring expedition. bulletin of the american museum of natural history 22, 131–134. https://doi.org/10.34194/geusb.v44.5342 http://www.geusbulletin.org https://dx.doi.org/10.17850/njg97-1-02 https://dx.doi.org/10.34194/geusm.v8.4526 https://dx.doi.org/10.34194/geusm.v8.4526 https://gallica.bnf.fr/ark:/12148/bpt6k4807624c/f11.item.texteimage.zoom https://gallica.bnf.fr/ark:/12148/bpt6k4807624c/f11.item.texteimage.zoom https://dx.doi.org/10.34194/geusb.v5.4800 https://dx.doi.org/10.34194/geusb.v5.4800 geological survey of denmark and greenland bulletin 11, 163-178 163 structural analysis of the northern nagssugtoqidian orogen, west greenland: an example of complex tectonic patterns in reworked high-grade metamorphic terrains stanislaw mazur, sandra piazolo and g. ian alsop structural analysis of the deeply eroded northern flank of the palaeoproterozoic nagssugtoqidian orogen shows marked regional variations in both the orientation and type of fabrics, as is characteristic of precambrian high-grade terrains subjected to polyphase deformation. here we investigate the relationship between strain, metamorphic grade, and the resulting structural patterns. the study area south of aasiaat in west greenland consists of amphiboliteto granulite-grade archaean orthogneisses and relatively thin supracrustal units. the regional foliation displays a wsw–ene to sw–ne strike associated with steep to moderate dips towards the wnw or sse. lineation trends are wsw–ene and generally plunge gently towards the wsw. mesoscopic fold hinges are usually colinear with the regional lineation. a systematic change in the plunge of lineations occurs across the south-western part of the study area. towards the south, the lineation plunge progressively increases, despite the generally uniform strike of foliation. this southward increase of lineation pitch is typically associated with the transition from l > s or l = s shape fabrics in rocks characterised by a low pitch, to s > l or s fabrics in the zone of moderate to high pitch. the structural patterns point to subdivision of the study area into a southern domain mostly characterised by s or s > l shape fabrics and a moderate to high angle of lineation pitch, and a northern domain showing l > s or l = s fabrics and low angles of lineation pitch. this subdivision corresponds well with the map scale boundary between granulite facies rocks in the south and amphibolite facies rocks farther north. the observed structural pattern may be explained by two alternative tectonic models: (1) northward indentation of the previously cooled granulite block into the rheologically weaker amphibolite domain, and (2) strain partitioning within a mid-crustal transpression zone. in model 2 the northern domain represents a localised zone dominated by strike-slip kinematics, whereas the southern domain shows evidence of mostly coaxial shortening. recent geochronology supports the indentator model in spite of limited available data. despite the details and structural complexities of the two tectonic models, the granulite and amphibolite facies domains seem to form autochthonous segments of a crustal section linked by a transitional zone that was only reactivated and reworked during indentation or transpression. the nagssugtoqidian compression was effectively transferred across this zone towards the northern amphibolite domain that suffered penetrative deformation during the palaeoproterozoic event. the n–s shortening was accommodated through folding, indentation and/or strike-slip displacements, rather than by thrusting and folding as seen south of the study area. keywords: deformation, gis, nagssugtoqidian orogen, transpression, indentation tectonics, west greenland ____________________________________________________________________________________________________________________________________________________________________________________ s.m., institute of geological sciences, university of wroclaw, maxa borna 9, 50-204 wroclaw, poland. e-mail: smazur@ing.uni.wroc.pl s.p., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. present address: department of geolog y and geochemistry, stockholm university, 10691 stockholm, sweden. g.i.a., school of geography and geosciences, university of st. andrews, fife ky16 9al, uk. © geus, 2006. geological survey of denmark and greenland bulletin 11, 163–178. available at: www.geus.dk/publications/bull 164 50 km 69° 67° 54° 50° disko bugt b a n ag ss ug to qi di an o ro ge n inland ice nno sno cno xxxx 68°15' kangaatsiaq tunorsuaq 53° 53° 52°30' 52°30' 53°30' 53°30' 53° 52°30'53°30' qeqertarsuatsiaq 10 km naternaqnaternaq naternaq saqqarputataneq ar fer s io r f ik amphibolite facie s naternaq 500 km surficial deposits quaternary basalt palaeogene amphibolite nagssugtoqidian orogen sisimiut charnockite arfersiorfik quartz diorite archaean orthogneiss, reworked metasedimentary rocks (mainly archaean) metasedimentary rocks (mainly palaeoproterozoic) granodioritic-granitic gneiss orthogneiss (unreworked) archaean craton orthogneiss, mainly tonalitic granitic gneiss metagabbro-anorthosite metamorphic transition zones quartzo-feldspathic metasediment mica schist and amphibolite quaternary marine deposits palaeogene dolerite dyke granulite fa cie s granulite fa cie s tra nsiti on zo ne tra nsiti on zo ne fig. 1. a: schematic geological map of the nagssugtoqidian orogen and adjacent foreland (modified from van gool et al. 2002b). outlined box shows location of the study area. sno, southern nagssugtoqidian orogen; cno, central nagssugtoqidian orogen; nno, northern nagssugtoqidian orogen. b: simplified geological map of most of the study area (modified from van gool et al. 2002a). 165 the deeply eroded northern flank of the palaeoproterozoic nagssugtoqidian orogen is exposed south of aasiaat in central west greenland (fig. 1). these rocks bear record of tectono-thermal processes that operated at mid-crustal levels in a collisional setting and controlled the distribution of strain and metamorphic facies. the area shows a complex structural pattern that varies significantly from south to north. the aim of this work is to describe the regional variation of structural elements and to investigate the relationship between strain, metamorphic grade and the orientation of deformational structures. this then leads to the consideration of two different tectonic models, which have been developed to account for the observed structural pattern: (1) an indentor model proposed originally by piazolo et al. (2004), and (2) a transpression zone model. the study area covers the kangaatsiaq geological map sheet at scale 1:100 000 (garde 2004), mapped in 2001– 2002 by the geological survey of denmark and greenland (geus; van gool et al. 2002b). the directional and fabric type structural data sets were analysed using geographic information systems (gis) based techniques that proved to be a powerful tool in the investigation of complex structural patterns. some of the structural data collected during this field campaign were presented by piazolo et al. (2004) to illustrate the application of gis in a multidisciplinary approach to survey high-grade terrains. our present study focuses on a more detailed analysis and interpretation of the structural relationships of the investigated area, a high-grade precambrian terrain affected by more than one deformation phase, and whose interpretation is not unequivocal. geological setting the study area covers over 3000 km2, and extends from 68°n–68°30′ n and 52°w–53°15′ w, between the fjord of ataneq in the south and the island of qeqertarsuatsiaq in the north (fig. 1). the area comprises the northern, c. 300 km wide exposure of the roughly e–w-trending nagssugtoqidian orogen (fig. 1). in the broadest sense, this tectonic belt resulted from a continent–continent collision between the archaean north atlantic craton to the south and an archaean continental mass to the north (e.g. kalsbeek et al. 1987; connelly et al. 2000; van gool et al. 2002a). the orogen is generally characterised by e–wtrending kilometre-scale folds and ene–wsw-trending linear belts which overprint an archaean fabric. on the basis of the grade of metamorphic reworking, ramberg (1949) and later marker et al. (1995) distinguished southern, central and northern segments of the nagssugtoqidian orogen (sno, cno and nno respectively; fig. 1). detailed structural investigations within the cno show that deformation in this area is dominated by thrust tectonics (manatschal et al. 1998; van gool et al. 1999). the investigated area lies within the nno and is composed of amphiboliteto granulite-grade archaean orthogneisses interlayered with relatively thin metasedimentary units (fig. 1). this region is transected by major fjord systems that allow data collection along well-exposed coastal sections. reconnaissance studies (noe-nygaard & ramberg 1961; henderson 1969; marker et al. 1995; kalsbeek & nutman 1996; mengel et al. 1998; connelly et al. 2000) provided initial information on the structural style of the study area. a comprehensive description of the structural pattern was recently presented by piazolo et al. (2004) and interpreted in terms of indentor tectonics with a rigid granulite-grade domain moving northwards into a rheologically weaker amphibolite facies domain. although quartzofeldspathic orthogneiss dominates in the investigated area, the overall map pattern is governed by discontinuous ne–sw-trending supracrustal belts (fig. 1). these 2–3 km thick sequences comprise several distinct lithological types: (a) monotonous, garnet-bearing quartzofeldspathic paragneiss locally containing subordinate mafic volcanic and metapelitic intercalations; (b) pelitic to semipelitic schist with or without garnet and sillimanite, including thin quartzofeldspathic layers and rarely quartzite; and (c) layered mafic to intermediate metavolcanic successions with calc-silicate bands and/or pods. previous studies have demonstrated that the metamorphic grade of the nno decreases northwards and is predominantly amphibolite facies, with granulite facies rocks preserved only in the south-western corner of the study area near the boundary with the cno (e.g. marker et al. 1995). the contact between these two facies is transitional over a distance of 10–12 km and forms a zone nearly parallel to the strike of the regional foliation (fig. 1). the granulite facies gneisses are typically pyroxene-bearing and enclose frequent melt pockets and cross-cutting veins. thermobarometric analyses point to a peak temperature of 800 ± 30°c at medium pressures of 6–7.5 kbar (piazolo et al. 2004). the amphibolite facies gneisses are lighter coloured and contain fewer biotite-bearing melt veins. they reveal peak metamorphism conditions of 650 ± 30°c at 4–5 kbar (piazolo et al. 2004). within the supracrustal amphibolites, crystallisation of amphibole, usually developed along foliation planes, indicates a syntectonic fluid flux and associated metamorphism. no relics of earlier granulite facies assemblages are preserved in these rocks. thrane & connelly (2006, this volume) carried out seve166 10 cm 50 cm 50 cm 30 cm 50 cm 50 cm a n s n s s n s n n s s n b c d e f fig. 2. examples of mesoscopic folds in the orthogneisses and metasediments of the study area. a: main foliation s 1 of orthogneiss developed parallel to the axial plane of isoclinal f 1 fold. b: lithological boundaries and originally cross-cutting basic dykes folded by f 1 isoclinal fold and showing extensive migmatisation. c: s 1 foliation folded into isoclinal folds during the same progressive d 1 event. d: s 1 foliation and its subsequent folding accompanied by pervasive migmatisation. e: s 1 foliation locally folded by f 3 folds of variable geometry with fold axes developed subparallel to the lineation. f: s 1 foliation of the orthogneiss folded by the f 3 fold with fold axes parallel to the lineation and s-directed asymmetry. n, north; s, south. 167 ral laser ablation and ion probe age determinations of zircon from within and adjacent to the present study area. deformation in the south-western part of the area is constrained by the emplacement age of 2748 ± 19 ma for a synkinematic granite which intrudes the orthogneiss. a palaeoproterozoic deposition age of c. 1950 ma was obtained from a metasediment within the naternaq supracrustal belt (fig. 1), and broad rims of zircons from an archaean sediment from kangersuneq yielded a metamorphic age of c. 1850 ma, suggesting that major nagssugtoqidian deformation and metamorphism occurred at around this time. thrane & connelly (2006, this volume) also obtained an age of 1837 ± 12 ma for a vertical, straight pegmatite north-east of kangaatsiaq trending 020° and displaying sinistral shear along its margins, that is thought to date a late phase of overall n–s-directed palaeoproterozoic shortening. characteristics of directional structures the main foliation (s 1 ) is axial planar to rare isoclinal f 1 folds that fold lithological boundaries as well as cross-cutting basic dykes (fig. 2a). these dykes have been rotated into parallelism with the foliation on the fold limbs (fig. 2b). the s 1 foliation was itself later folded into isoclinal folds, although this refolding may reflect the same progressive d 1 deformation event since no overprinting fabric is associated with it (fig. 2c). a characteristic feature of these folds is the broad parallelism of their axial planes to the regional foliation s 1 . a l 1 mineral lineation is developed on the foliation planes and is usually defined by a parallel alignment of amphibole crystals and/or elongated quartz-feldspar and biotite aggregates. the lineation is well developed in the amphibolite facies rocks but rather weak in the granulite facies gneisses. in the transition zone between the amphibolite and granulite domains (fig. 1b), no mutually cross-cutting mineral lineations were detected and no evidence for fabric superimposition was observed. the l 1 lineation is only rarely associated with kinematic indicators that are commonly symmetric and must have resulted from coaxial strain and/or a finite strain combining the effects of several strain increments. asymmetric fabrics have been observed only in zones of steeply dipping foliation, and typically indicate a sinistral rotational shear component in the present-day coordinates (fig. 3). restoration of the steep foliation attitude to more gentle regional dips would result in the same indicators implying a top-to-the-west or wsw sense of shear. the s 1 foliation and its subsequent folding during the presumed progressive d 1 event, were accompanied by a long-lasting pervasive migmatisation. this is demonstrated by the common occurrence of migmatite layers or patches that are variably deformed and show mutually cross-cutting relationships. some of them are parallel to the main foliation s1 (fig. 2d) whereas others define discordant veins or dykes oblique to the regional fabric. between these two end members are a range of cross-cutting veins that are deformed and reoriented to varying degrees. in the metasedimentary rocks of the naternaq (lersletten) area, the s1 regional foliation is refolded by f2 folds characterised by steep to subvertical fold hinges. these folds are developed at kilometreto centimetre-scale and, in a few cases, map scale f2 folds can be seen refolding f1 (a.a. garde & j.a. hollis, personal communication 2003). f2 folds are found exclusively within the metasedimentary belts and at their contacts with the adjacent gneisses. at a mesoscopic scale, they are represented by folds plunging steeply towards the se and in few cases towards the north. a moderate to strong, se-plunging mineral lineation is associated with the hinges of f2 folds, locally deviating a b 20 cm 5 cm wsw ene wsw ene fig. 3. examples of sinistral (top-to-the wsw) kinematic indicators in the orthogneisses of the study area. the lineation is plunging towards wsw on north-dipping foliation planes. a: sheared, asymmetric amphibolite enclave. b: sigmoidal k-feldspar porphyroclast. 168 from its regional trend although it remains the only lineation present. an s 2 cleavage axial planar to f 2 is rarely, and only weakly, developed. beyond naternaq, the s 1 foliation is refolded by f 3 folds of variable, often complex geometry with gently plunging fold axes developed subparallel to the l 1 lineation (fig. 2e). the frequency of such folds increases northwards within the amphibolitegrade rocks. associated f 3 axial planes are locally marked by a subtle s 3 cleavage, accentuated by local mica aggregates and discrete joints which cut the main s 1 foliation. the distinction between f 2 and f 3 folds is primarily based on their different geometries since they both fold the regional fabric s 1 and are not associated with penetrative axial cleavages or intersection lineations. the hinges of f 2 folds are relatively steep and oblique to the regional lineation, whereas f 3 axes are gently inclined and run parallel to the lineation l 1 . the f 3 mesoscopic folds are frequently asymmetric with fairly uniform sse vergence in the area south and west of kangaatsiaq (fig. 2f). orientation of directional structures on the map scale, the regional foliation displays a wsw– ene to sw–ne strike associated with steep to moderate dips towards the nnw or sse (fig. 4). shallow-dipping foliations (< 30°) are rare and randomly distributed throughout the study area. their variable directions suggest that they are associated with the hinge zones of f3 folds developed at different scales (fig. 4). moderately and steeply dipping foliations show a distinctly discrete grouping within the investigated area. the former predominate in the south-eastern corner of the area, corresponding with the granulite-grade block, whereas the steeply dipping foliations are developed in its central part, forming a wide belt along tunorsuaq (fig. 4). this belt coincides with a transition zone between the granulite and amphibolite facies domains and partly with the south-eastern margin of the latter (see figs 1, 4). the strike of foliation dipping steeper than 30° remains fairly consistent throughout the area, while the dip direction varies only in the case of sub68°30' 68° tunorsuaq ataneq naternaq 53°53° 52°30'52°30' 52°52°53°30' 53° 52°30' 52° ab tz tz gb 68°15' 10 km ≤ 30° foliation 30–60° > 60° fig. 4. representation of foliation trends and dip directions in the study area. gb, granulite block; tz, transition zone; ab, amphibolite block. 169 vertical planes (> 60°). on stereoplots, the poles to foliation are scattered along a regular girdle produced by the f3 folds (fig. 5). in the southern granulite facies block the vast majority of foliation measurements cluster in one maximum, suggesting that late folding is insignificant or absent in this area (fig. 5). this maximum corresponds to the foliation dipping moderately to steeply towards the nnw, while the axis of the foliation girdle plunges gently to the wsw at 20°. in the northern, amphibolite facies block measurements are more evenly distributed along the foliation girdle, providing evidence for the regional importance of f3 folding (fig. 5). in contrast to the southern block, the stereographic girdle axis is almost subhorizontal. although the foliation patterns in the southern and northern blocks are fairly similar, a striking difference occurs within the latter between the rheologically competent orthogneisses and relatively incompetent supracrustal formations (fig. 6). the orthogneisses reveal a regular girdle perfectly controlled by a cylindrical geometry of the f3 folds. in contrast, the less competent supracrustal rocks display a pronounced foliation scatter, reflecting a complex interference between the effects of the (possibly noncylindrical) f2 and f3 folding. the lineation trends wsw–ene over the whole study area, and frequently plunges gently towards the wsw (fig. 7). shallow lineations (< 15°) are concentrated in the northern and central parts of the investigated area (fig. 7), and commonly coincide with the subvertical foliation within the amphibolite facies block and the associated transition zone towards the granulite facies block. conversely, lineations plunging steeper than 15° group largely in the south-eastern corner of the area, corresponding with the granulite block, and show mostly wsw-directed plunges. mesoscopic f3 fold hinges are usually colinear with the regional lineation, and this is also the case at the larger scale since the lineation maxima are located near the pole of the foliation girdle on stereoplots (cf. fig. 5). in the amphibolite facies block, the lineation is shallow and relatively well grouped in the maximum representing the subhorizontal wsw–ene trend. a significant scatter occurs only in the naternaq area, which is displayed on the map by relatively steeply plunging lineations (fig. 7) and on stereoplots by data distribution along a small cira b c d northern block southern block n n n n n = 794 n = 616lineation lineation n = 794 n = 616foliation foliation lineation maximum 245/05 lineation maximum 245/05 245/05 f3 245/20 f3 f2 f2? 6% 4% 2% 0.5% 8% 6% 4% 2% 0.5% fig. 5. attitudes of foliation s 1 and stretching lineation l 1 in the northern amphibolite facies block (a, c), and southern granulite facies block (b, d). the positions of the poles to the foliation girdles and maximum of lineation measurements are indicated in stereoplots a, b and c, d, respectively. 170 cle (fig. 5). a dispersal of lineation measurements is produced in that area by steeply plunging f 2 fold hinges whose orientation corresponds to the centre of a small circle defined by the scatter of the lineation (fig. 5). two different small circles can be delineated on the stereoplot (fig. 5) based on the lineation scatter. one is centred at the orientation of f 2 hinges steeply inclined to the se that are relatively frequent as mesoscopic structures. the second is developed around the nearly vertical north-plunging f 2 hinges rarely found in the outcrops but probably important at the map scale. the lineation scatter induced by the f 2 folding is very clear in the supracrustal rocks, whereas it is almost absent in the orthogneisses (fig. 6). this relationship is consistent with the field observation that the f 2 folds are developed almost exclusively in the metasediments. in the southern block the lineation is slightly steeper than in the north and shows a mean plunge of c. 20°. its maximum is more diffuse than in the amphibolite-grade block and more lineations are relatively steep (> 15°). nevertheless, the average lineation trend defined by the position of maxima on stereoplots is exactly the same for the amphibolite and granulite facies domains (fig. 5). in order to better understand the geometric relationships between planar and linear fabric elements during deformation, they may be directly compared on fabric topology plots in terms of fabric trends and lineation pitch (see alsop & holdsworth 2004 for a review). the angle of pitch may be defined as the angle that a line makes with the strike of a surface, when measured within that plane (fig. 8a). a significant variation in the angle of lineation pitch on the regional foliation surface is observed (fig. 8b); this may be caused by two independent factors: (1) variable plunge of the lineation, and (2) variable dip direction of the foliation. the latter feature seems to be a consequence of local folding, since a high pitch angle (> 45°) is most characteristic for sw-dipping foliations (fig. 9) that represent the hinges of f 3 folds (fig. 5). on the other hand, a majority of measurements correspond to foliation dipping to the nw or se that reveals a low or moderate pitch angle (fig. 9). such a foliation pattern is consistent with the regional attitude of foliation inclined towards the nw and only locally reoriented on limbs of the f 3 folds. since the f 3 folds are only of minor significance in the southern part of the study area characterised by higher pitch values (fig. 8), and the majority of lineations were measured on steep to subvertical foliation sura b c d orthogneisses supracrustal rocks n n n n n = 554 n = 237 lineation lineation n = 554 n =237 foliation foliation lineation maximum 245/05 lineation maximum 245/05 6% 4% 2% 0.5% 8% 6% 4% 2% 0.5% 245/05 f2 f2? fig. 6. attitudes of foliation (s 1 ) and stretching lineation (l 1 ) in the orthogneisses (a, c) and supracrustal rocks (b, d) of the northern amphibolite facies block. the position of the pole to the foliation girdle and maximum of lineation measurements are indicated in stereoplots a and c, d, respectively. 171 faces, the broad scatter of the pitch angles can be attributed to variations in the lineation plunge (fig. 10). only a minority of measurements plot away from the line that indicates the similarity of pitch and plunge angles (fig. 10). thus, the change of foliation strike plays a less important role in the distribution of pitch angles. the low pitch angle (< 15°) is characteristic of the central and northern parts of the study area (fig. 8), corresponding to the amphibolite facies block and the transition zone to the granulite facies block. this is the area that is additionally characterised by the steep foliations and shallow lineation plunge. the higher pitch angles (exceeding 15°) are more common in the south-east corner of the study area within the granulite facies gneisses, where they are related to relatively steep lineations occurring on the moderately to steeply inclined foliation. a systematic change in the plunge of lineations can be observed across the south-western part of the study area (fig. 7). the steep sw–ne-striking foliation around tunorsuaq is associated with the gently plunging lineation that defines a low angle of pitch on the foliation. towards the south, however, the plunge of lineation becomes progressively steeper despite the generally uniform strike of foliation (fig. 4). a pitch versus lineation trend diagram (fig. 11) demonstrates that the increase in pitch is unrelated to the change of lineation trend. this means that lineations are typically not reoriented on fold limbs, and that folds, if present, are mostly colinear with lineation. furthermore, in the north-western part of the area the lineations plunge gently to the sw and ne to define a series of culminations and depressions, that are clearly illustrated by the opposing plunges on the south side of tunorsuaq and the outer islands to the west-south-west (fig. 7). this sinuosity of lineations and associated fold hinges defines a large-scale whaleback pattern consistent with a dominantly subhorizontal and approximately n–s contractional strain (cf. piazolo et al. 2004). the finite planar (s) and linear (l) shape fabrics within a high-strain rock may be qualitatively described (flinn 1978). consequently, the relative dominance of these re68°30' 68° tunorsuaq ataneq naternaq 53°53° 52°30'52°30' 52°52°53°30' 53° 52°30' 52° ab tz tz gb 68°15' 10 km < 15° lineation 15–35° > 35° fig. 7. representation of lineation trends and plunge directions of the study area. gb, granulite block; tz, transition zone; ab, amphibolite block. 172 68°30' 68° tunorsuaq ataneq naternaq 53°53° 52°30'52°30' 52°52°53°30' 53° 52°30' 52° ab tz tz gb 68°15' 10 km ≤ 15° pitch 15–40° > 40° b strik e dip angle lin ea tio n pitch angle dip foliatio n a 0 ne 90° se 180° dip direction 270° 360°nwsw 90° 60° 30° 0 pi tc h an gl e fig. 8. pitch data. a: schematic diagram illustrating the pitch of a lineation. b: representation of lineation pitch angles of the study area. gb, granulite block; tz, transition zone; ab, amphibolite block. fig. 9. plot of lineation pitch angle vs. dip direction of foliation, showing the change of pitch as a consequence of local folding. 173 spective components enables a distinction to be made between fabrics that are foliation dominated (s tectonite), lineation dominated (l tectonite) or contain a combination of foliation and lineation (sl tectonite). the pattern shown by the spatial distribution of fabric types in the study area is clearly differentiated. the north-western part of the area is dominated by ls and l > s fabric types whereas the south-eastern part reveals a vast preponderance of s or s > l types (fig. 12). the southward increase of lineation pitch is typically associated with the transition from l > s or ls shape fabrics in rocks characterised by a low pitch, to s > l or s fabrics in the zone of moderate to high pitch. constrictional fabrics, i.e. l and l > s fabrics, are typical of rocks with shallowly plunging lineations on the steep foliation belonging to the amphibolite facies block and the transition zone to the granulite facies block. the latter is dominated by the flattening fabric types, i.e. s and s > l fabrics, which are common in rocks containing steeply plunging lineations. discussion three main observations can be made concerning the data presented above: (1) structural elements, i.e. lineations, foliations, folding and fabric type, vary with lithology, (2) two broad domains with different structural patterns can be distinguished, and (3) the structural style in the nno is markedly different to that previously described from the cno, e.g. by manatschal et al. (1998) and van gool et al. (1999). the observed scarcity of mineral lineations in the gran90° 60° 30° 0 pi tc h an gl e 0 30° lineation plunge 60° 90° pitch an gle = plunge angle 0 ne 90° se 180° lineation trend 270° 360°nwsw 90° 60° 30° 0 pi tc h an gl e 0 ne 90° se 180° lineation trend 270° nwsw fig. 10. plot of lineation pitch angle vs. lineation plunge, showing the contribution of changing lineation plunge to the total variation of pitch. measurements located on the straight line connecting the lower left and upper right corners of the plot have been taken on vertical foliation planes. fig. 11. plot of lineation pitch angle vs. lineation trend, showing the variation of pitch despite the uniform trend of lineation. 174 ulite facies gneisses could be attributed to the absence of minerals with a high shape aspect ratio. if such minerals are lacking, lineations of the type that forms by rigid rotation or due to growth parallel to the maximum extension direction or transport direction can hardly develop. nevertheless, aggregate lineations that form by dynamic recrystallisation and material transfer should be expected in the granulite facies gneisses. since even these lineations are lacking in the southern block, it seems that the scarcity of linear structures in this area is not only a function of lithology but also reflects specific features of deformation history. in the northern block, a significant strain partitioning between orthogneisses and supracrustal formations is indicated by the effects of f2 folding in metasedimentary rocks and the resultant scatter of foliation and lineations in these rocks (fig. 6). this points to the concentration of deformation in rheologically weak metasedimentary belts that accommodate the bulk of strain induced during f2 folding. the competent orthogneisses experienced only subsequent f3 folding, when they were deformed together with associated less competent supracrustal rocks. at the western limb of the naternaq supracrustal belt, there is a structural discordance with another e–w-trending supracrustal unit farther to the west. this structural discordance may suggest the presence of some kind of ‘stockwerke tectonics’ (wegmann 1935), in addition to the inferred strain partitioning between the orthogneisses and metasedimentary rocks. the application of the ‘stockwerke’ model to the naternaq belt itself remains an intriguing problem that cannot be resolved at present, due to the lack of sufficient information on the time relationship between the structural discordance and the main phase(s) of folding and metamorphism. disregarding the local complexity at naternaq, the general structural pattern described in this paper allows subdivision of the study area into two main domains: (1) a southern domain mostly characterised by s or s > l shape fabrics and a moderate to high angle of pitch, and (2) a northern domain showing l > s or ls fabrics and low 68°30' 68° tunorsuaq ataneq naternaq 53°53° 52°30'52°30' 52°52°53°30' 53° 52°30' 52° ab tz tz gb 68°15' l or l > s fabric type l = s s or s > l 10 km fig. 12. spatial distribution of fabric types (terminology af flinn 1978). gb, granulite block; tz, transition zone; ab, amphibolite block. 175 angles of lineation pitch. this subdivision compares well with the map scale variation of the metamorphic grade from granulite facies in the south to amphibolite facies further north. the contact between these two facies is transitional and forms a gently curved boundary that is subparallel to the strike of the regional foliation (piazolo et al. 2004). this division of the study area into two contrasting blocks, indicated by the structural data and variation of metamorphic grade, can be explained by two models, namely (1) an indentor model, and (2) a transpressive deformation model, the relevance of which is briefly discussed below. the indentor model, originally developed by piazolo et al. (2004), invokes a twofold deformation history. during the first event the southern block was subjected to a coaxial flattening under granulite facies conditions (fig. 13b – time a). a fairly uniform strain developed at that time and is manifested by prevailing s or s > l shape fabrics. during the second event (fig. 13b – time b) the previously cooled granulite grade block acted as an indentor, with the amphibolite facies domain being plastered and moulded around the rigid block. the structural pattern and mineral assemblages of the southern domain were only modified in a transition zone that experienced retrogression to the amphibolite facies. at the same time, the amphibolite grade northern domain was subjected to mostly coaxial strain, including an important constrictional component. the resultant structural grain in the amphibolite facies rocks mimics the geometry shown by the boundaries of the southern block. the alternative model (fig. 13c) explains the observed structural relationships in terms of strain partitioning within a transpression zone (for a review of transpression see holdsworth et al. 2002, and references therein). in this model, the different structural patterns documented in the northern and southern domains were produced during a single deformation event. consequently, the model implies a continuity of the structural grain across the transitional zone as well as a gradual change in the orientation of directional structures, fabric type and kinematics of strain. the southern domain was mostly subjected to coaxial flattening, resulting in s or s > l shape fabrics with a moderately plunging lineation. the northern domain was deformed in a wrench-dominated regime, characterised by a constrictional or plane strain with a significant rotational strike-slip component. hence, the northern block reveals l > s and ls shape fabrics and shallowly plunging lineations. a sinistral sense of displacement assumed in the model is consistent with rare observations in the field of kinematic indicators with the appropriate asymmetry. the indentor model explains the arcuate swing of the structural grain around the granulite facies block, the map scale geometry of which may appear to control the foliation and lineation patterns within the amphibolite facies block further north. a serious weakness of the indentor model is the lack of evidence for fabric overprinting in the boundary zone between the two domains despite the assumed twofold deformation. however, high strain in the transitional zone could account for the obliteration of interference patterns and the apparent continuity of structural grain. the transpression zone model accounts for the presence of steep (f 2 ) fold hinges in the naternaq supracrustal belt that are expected to develop within the wrenchdominated part of the transpressive system. the model also explains the contrast of structural style between the northern and southern domain without a detectable tectonic boundary or an overlap of structural patterns. consequently, it is consistent with an apparent continuous transition linking the structural patterns within the granulite and amphibolite blocks. the somewhat steeper plunge of lineations within the granulite facies domain may indicate a greater component of coaxial strain in that area. however, the potential strain partitioning, as revealed by the lineation pattern, is relatively weak for a transpression zone. the partitioning of strain that takes place between the orthogneisses and supracrustal rocks within the northern amphibolite facies block is consistent with both of the discussed models. in addition, both models are consistent with a dominantly subhorizontal and approximately n–s contractional strain in the northern block causing the small-scale porpoising and large-scale whalebacking of lineations and associated fold hinges. thus, even a detailed structural analysis does not allow us to determine unequivocally which of the two models is the more appropriate. nevertheless, the presented structural models can be indirectly verified by geochronological data, which constrain the age of peak metamorphism and associated deformation in the northern amphibolite facies block as palaeoproterozoic (thrane & connelly 2006, this volume). furthermore, the late archaean crystallisation age of syndeformational granites emplaced at saqqarput (fig. 1) and at the southern margin of the nno (connelly & mengel 2000; thrane & connelly op. cit.), points to a lack of significant deformation in the southern block from the late archaean onwards. this is in conflict with the coincidence of deformation between the northern and southern blocks required by the transpression model and, thus, supports the indentor model. this is corroborated by the occurrence of undeformed but metamorphosed mafic dykes of likely palaeoproterozoic origin emplaced at the southern margin of the investigated area (glassley & sørensen 1980; árting 2004). 176 ii iv i iii i iii v vi ii iv i ii v v ii iii iv vi ii i i iii iv vi i ii iii iv plan view granulite facies v vi i iii ii iv i iii iv i ii v ii i iii granulite facies amphibolite facies plan view v vi i ii iii iv v vi ii granulite facies amphibolite facies plan view v ii i ii ii v a schematic trace of lineation schematic trace of folds principal stress – σ1 response to stress model i (piazolo et al. 2004). observed fabrics developed during two deformational events model ii (discussed in this paper). fabrics developed during a single transpressional event time a, at granulite facies conditions time b, at amphibolite facies conditions c b pla ne st ra in a xi al r at io o n x y p la ne flattening constriction axial ratio on yz plane s or s > l fabric sl fabric boundary of deformation transition zone fig. 13. schematic indentor and transpression zone models with expected fabric types and fold patterns. a: flinn graph with schematic qualitative descriptions of the finite planar (s) and linear (l) shape fabrics. b: indentor model. c: transpression zone model. 177 the f 3 folding apparently postdates the juxtaposition of the granulite and amphibolite blocks and had a relatively minor influence on their contact zone. the northward increase in intensity of f 3 folding is readily explained by the rheological weakness of the amphibolite facies domain during cooling. the origin of f 3 folds provides evidence for prolonged, approximately n–s-directed compression, the age of which is roughly constrained by the previously mentioned 1837 ± 12 ma pegmatite (thrane & connelly 2006, this volume). our structural data also show that compressional stress related to the growth of the nagssugtoqidian orogen was effectively transferred across the lower crust, and that the amphibolite facies domain south of aasiaat was subjected to the penetrative palaeoproterozoic deformation. therefore, this area represents an integral part of the nagssugtoqidian orogen despite the obvious paucity of palaeoproterozoic crustal components. in contrast to the southern part of the orogen, the overall n–s shortening induced by the nagssugtoqidian collision was accommodated in the study area through indentation and folding, in contrast to the central part of the orogen that displays significant thrust tectonics. concluding remarks the indentor and transpression-driven tectonic models discussed in this paper share a number of features which shed a new light on the evolution of the northernmost nagssugtoqidian orogen. the granulite and amphibolite facies blocks distinguished in the study area appear to be (par)autochthonous, and their boundary is only reactivated and reworked during indentation or transpression. the area studied shows excellent examples of precambrian deformation that are characterised by significant strain partitioning into less competent metasedimentary rocks. the presence or absence of lineations may be influenced significantly by lithology and metamorphic grade. complex structural patterns may not always conclusively reveal the structural history on their own, hence geochronological data are essential in distinguishing between different deformation models such as transpression or indentation. in addition, it is apparent that the tectonic styles of the lower to middle crust can be highly variable within the same orogen. deformation may vary from thrust-dominated in the central portion to folding associated with indentation or transpression on its flanks. acknowledgements critical reviews by c.r.l. friend and an anonymous reviewer as well as constructive comments by a.a. garde helped us to improve the manuscript, and are gratefully acknowledged. structural information acquired from other members of the geus field parties in 2001–2002 significantly contributed to the data base used in this paper. references alsop, g.i. & holdsworth, r.e. 2004: the geometry and topology of natural sheath folds: a new tool for structural analysis. journal of structural geology 26, 1561–1589. árting, u.e. 2004: a petrological study of basic dykes and sills of assumed palaeoproterozoic age in central western greenland, 122 pp., two appendices. unpublished m.sc. thesis, university of copenhagen, denmark. connelly, j.n. & mengel, f.c. 2000: evolution of archean components in the paleoproterozoic nagssugtoqidian orogen, west greenland. geological society of america bulletin 112, 747–763. connelly, j.n., van gool, j.a.m. & mengel, f.c. 2000: temporal evolution of a deeply eroded orogen: the nagssugtoqidian orogen, west greenland. canadian journal of earth sciences 37, 1121–1142. flinn, d. 1978: construction and computation of three-dimensional progressive deformations. journal of the geological society (london) 135, 291–305. garde, a.a. 2004: geological map of greenland, 1:100 000, kangaatsiaq 68 v.1 syd. copenhagen: geological survey of denmark and greenland. glassley, w.e. & sørensen, k. 1980: constant ps-t amphibolite to granulite facies transition in agto (west greenland) metadolerites: implications and applications. journal of petrology 21, 69–105. henderson, g. 1969: the precambrian rocks of the egedesminde–christianshåb area, west greenland. rapport grønlands geologiske undersøgelse 23, 37 pp. holdsworth, r.e., tavarnelli, e., clegg, p., pinheiro, r.v.l., jones, r.r. & mccaffrey, k.j.w. 2002: domainal deformation patterns and strain partitioning during transpression: an example from the southern uplands terrane, scotland. journal of the geological society (london) 159, 401–415. kalsbeek, f. & nutman, a.p. 1996: anatomy of the early proterozoic nagssugtoqidian orogen, west greenland, explored by reconnaissance shrimp u-pb dating. geology 24, 515–518. kalsbeek, f., pidgeon, r.t. & taylor, p.n. 1987: nagssugtoqidian mobile belt of west greenland: a cryptic 1850 ma suture between two archaean continents – chemical and isotopic evidence. earth and planetary science letters 85, 365–385. manatschal, g., ulfbeck, d. & van gool, j.[a.m.] 1998: change from thrusting to syncollisional extension at a mid-crustal level: an example from the palaeoproterozoic nagssugtoqidian orogen (west greenland). canadian journal of earth sciences 35, 802–819. 178 marker, m., mengel, f., van gool, j.[a.m.] & field party 1995: evolution of the palaeoproterozoic nagssugtoqidian orogen: dlc investigations in west greenland. rapport grønlands geologiske undersøgelse 165, 100–105. mengel, f., van gool, j.a.m., krogstad, e. & the 1997 field crew 1998: archaean and palaeoproterozoic orogenic processes: danish lithosphere centre studies of the nagssugtoqidian orogen, west greenland. geology of greenland survey bulletin 180, 100–110. noe-nygaard, a. & ramberg, h. 1961: geological reconnaissance map of the country between latitudes 69°n and 63°45´n, west greenland, 1:500 000. geological map grønlands geologiske undersøgelse 1, 9 pp., 2 maps. (also meddelelser om grønland 123(5).) piazolo, s., alsop, g.i., van gool, j.[a.m.] & nielsen, b.m. 2004: using gis to unravel high strain patterns in high grade terranes: a case study of indentor tectonics from west greenland. in: alsop, g.i. et al. (eds): flow processes in faults and shear zones. geological society special publication (london) 224, 63–78. ramberg, h. 1949: on the petrogenesis of the gneiss complexes between sukkertoppen and christianshaab, west greenland. meddelelser fra dansk geologisk forening 11, 312–327. thrane, k. & connelly, j.n. 2006: zircon geochronology from the kangaatsiaq–qasigiannguit region, the northern part of the 1.9– 1.8 ga nagssugtoqidian orogen, west greenland. in: garde, a.a. & kalsbeek, f. (eds): precambrian crustal evolution and cretaceous– palaeogene faulting in west greenland. geological survey of denmark and greenland bulletin 11, 87–99 (this volume). van gool, j.a.m., kriegsman, l., marker, m. & nichols, g.t. 1999: thrust stacking in the inner nordre strømfjord area, west greenland: significance for the tectonic evolution of the palaeoproterozoic nagssugtoqidian orogen. precambrian research 93, 71–86. van gool, j.a.m. et al. 2002a: precambrian geology of the northern nagssugtoqidian orogen, west greenland: mapping in the kangaatsiaq area. geology of greenland survey bulletin 191, 13–23. van gool, j.a.m., connelly, j.n., marker, m. & mengel, f.c. 2002b: the nagssugtoqidian orogen of west greenland: tectonic evolution and regional correlations from a west greenland perspective. canadian journal of earth sciences 39, 665–686. wegmann, c.e. 1935: zur deutung der migmatite. geologische rundschau 26, 305–350. __________________________________________________________________________________________________________________________________________________________________ manuscript received 17 june 2004; revision accepted 14 june 2005 geological survey of denmark and greenland bulletin 35, 2016, 99-102 99© 2016 geus. geological survey of denmark and greenland bulletin 35, 99–102. open access: www.geus.dk/publications/bull since 2008 the geological survey of denmark and greenland (geus) has used handheld devices (personal digital assistents pdas) for collecting field data in a digital format. since pdas are becoming obsolete and new device technology with improved functionality is available, it was decided to develop an android-based application (app) that can be used by many mobile telephones and to test this on different devices during field work in greenland. the main objectives of the system are to provide field geologists working in remote areas without internet access with a quick and efficient way of: (1) recording information on a geological locality in digital format, (2) displaying existing digital geodata on maps, (3) securing a consistent way of reporting data and (4) transferring data quickly and efficiently to the central databases and gis environments once the field teams return from their field work. with the handheld device and the app (the system) the user collects data much in the same way as it was previously done using the modified ganfeld software with the pdas (schlatter et al. 2010), but with improved functions and a more user-friendly interface. the development of the system was guided by the following prerequisites: • the device has to be relatively robust, lightweight, small and easy to handle, • the device has to be easily replaceable and relatively cheap, • the device includes an internal gps and camera of sufficient quality with low power consumption; recharging must be with solar panels, or a mobile generator, • the device must have backup, both internal (sd memory card) and external (laptop) during field work, the app must be independent of mobile networks and must work offline, • the app must store data in a well-structured format allowing for easy transfer to other gis environments, • the system should allow for both data collection and display in the field, and editing of data in the field camp, • new data as well as existing geodata and maps (topographic data, geological maps, geochemical, geophysical, etc.) should be displayed together, and easily imported into the app, • the system should have a simple, intuitive, user-friendly interface that is easy to operate in the field under difficult conditions. the decision to develop an app for the android platform was based on (1) the wide variety of relatively cheap android devices available in many different sizes (including several robust versions), (2) the java development language runs on a windows platform, the main development platform used at geus, and (3) the possibility to store data on a sd removable memory card that can easily be transferred to a laptop or pc in the field camps. the afieldwork app the app was designed to require as little typing as possible. this was done to make digital capture of field data as easy and efficient as possible. the app is database-driven meaning that data have to be entered in a structured and consistent fashion. although some of the data entry is compulsory and is entered via scrollable, drop-down lists of pre-defined and fixed content, other information is optional and includes free-text entry. the mandatory entry points and the predefinition of selection lists also induce the field geologists to describe and classify localities and geological features in a consistent way. the first version of the app was a prototype allowing for sufficient collection of information about localities, lithology, samples, structures and photographs. the system was further developed and refined based on field tests of the prototype in greenland. online and offline maps two different map types are implemented in the app. the user can choose to use (1) online maps from openstreetmap (if connected to the internet), or (2) pre-loaded openstreetmap or a series of pre-loaded maps stored in a database. the last option is by far the most advanced and requires considerable preparation, but provides the most flexibility including custom-drawn maps at different zoom levels (scales) as well as various overlays of other data types (fig. 1). afieldwork – an android app for offline recording of geological information and data display martin hansen, matti nelleman petersen, thomas f. kokfelt and bo møller stensgaard 100100 all map types show and create geological localities as symbols. as long as the gps is on, a colour-coded indicator shows each locality and the type of information which is attached to the locality. the user can make drawings and queries about spatial data, such as contours, trend lines and former localities. from the map a click on any locality or point will display attached information which can then be edited. while the map is displayed on the screen the user’s position and walking direction will be shown. it is also possible to create points on the map that are not related to geological observations, a function useful for planning reconnaissance missions or drop-offs in the field by helicopter or boat. the user can make distance measurements on the maps and quickly obtain the coordinates of any point. how to enter data all data entries are geo-referenced via geographical co-ordinates taken from an internal gps in the android device together with information about its accuracy. the sample numbering follows geus’ sample numbering system. the user may prompt manual text entries, for e.g. longer descriptions of localities, rocks and samples. for example, the data entry point, or menu, for samples will not appear before the user selects the ‘sample button’. this ensures a smooth and efficient data entry without interference of unnecessary information and menus. the selection lists are in many cases designed so that the most frequently used entries are listed first and can be tailor-made before field use for special requirements. ‘localities’. a locality is created by tapping the ‘new locality’ button on the start-up screen of the afieldwork app (fig. 2). the user is then prompted for a description of the locality (not mandatory). the co-ordinates are displayed as decimal degree, degree decimal minutes or converted to utm coordinates, along with their gps accuracy and time of creation. the locality is automatically stored and provided with a unique locality id which is composed of the year, the unique geus initials of the geologist and the number of the locality, e.g. 14smw003 for the third locality of the geologist smw in the year 2014. from the locality screen it is possible to add free text notes or to open entry points for ‘earth materials’, ‘samples’, ‘structure’ or ‘photos’. each entry point has been colour-coded to facilitate identification (fig. 2). ‘earth materials’ (lithology). the entry point ‘earth materials’ (fig. 3) refers to information about the lithologies found at a given locality. this information is entered from a predefined, hierarchical system of ‘rock class’, ‘rock type’ and subsequently ‘rock name’ allowing a narrowing down to a few items with a few taps on the screen. the classification of the ‘earth materials’ (the rocks) follows classification schemes suggested by the international union of geological sciences (iugs). it is possible to omit a rock name at ‘earth materials’. extra information, such as colour, metamorphic grade, mineralogy (as ‘common’, ‘alterations’ and ‘ore’ minerals) and fossils can be entered. this information is mostly entered via predefined selection lists. once the ‘earth materials’ data are entered it can be saved, not only to the locality, but also to a quick list (‘quiklst’). this makes it easy to add information from this quick list to additional localities with the same ‘earth material’ information. ‘samples’. at the entry point ‘sample’ the user can number and register information on particular rock samples collected in the field. the rock samples must be connected to an fig. 1. a: geological (1:500 000) map with layers of topographic contours and various locality point symbols. b: the same geographical area with an overlay of the analytical signal of the total magnetic intensity on top of the geological map. a b 101 ‘earth material’ entry and cannot be saved without this critical information. the ‘earth material’ entry can either be an existing one or can be created when the ‘sample’ entry point is selected. the sample is automatically assigned a sample number (by default one increment above the previous sample number). a sample type and a sample purpose can be selected from drop-down lists and a free text can also be entered. structural measurements. at the entry point ‘structure’ the user can digitally capture information on a structural measurement. like rock samples, structural observations must be connected to an ‘earth material’. structural class (linear or planar), type (e.g. fault, shear zone for a planar class) and detail (e.g. generation for a foliation) can be selected from drop-down lists. measurements made by a hand-held compass can be entered as numbers using a keyboard or by two sliding bars. ‘photos’. photographs can be taken with the internal camera, in which case there is no need for data entry besides an optional free text description of the photograph. if an external camera is used, the user can enter the first and last number of the photographs taken at a given locality. it is also possible to enter free text for the photograph taken with an external camera. sketches and annotations can be added to the photographs taken by the internal camera. they will be saved as separate files leaving the original photograph intact. behind the screen all data are stored in an sqlite database used in most small digital devices. it is a relatively advanced database system supporting foreign keys, triggers and (with an extension) spatial data. the back-end database structure ensures wellstructured, compatible data and facilitates data transfer to fig. 2. the main screen of the app showing a list of localities with colour code added for easy reference: ‘earth materials’ (lithology), ‘samples’, ‘structure’ measurements or ‘photos’. fig. 3. the ‘earth material’ (lithology) data entry screen. 102102 other databases. because the predefined selection lists in afieldwork app are based on the database queries it is easy to extend the system or to configure it to suit the geological requirements of specific areas. part of the application (app) is based on open source software. the database handling is based on the android sqlite manger (asqlitemanger), the map is based on the mapsforge library (mapsforge) and the spatial functions on the spatialite library (spatialite). future development a future update of the app could be improved by adding a number of new features: • import of points from google earth, • built in documentation of all ‘earth materials’ and ‘structure’ classes, • an easy overview of all structural data by symbols, • the possibility to choose a location on the map and be guided there, • easy export of data for analysis and planning of field work in field camps, • use of air-pressure sensors to improve altitude determinations, • new information fields for registering the type of locality, e.g. geological locality, field camp, archaeological sites, emergency pick-up site, etc., • display of the paths taken during field work, • display geo-referenced raster images (e.g. images of geophysical, geochemical and remote sensing data) which will most likely be based on the rasterlite library (rasterlite). future versions of the app will enable the user to handle other sample types, e.g. stream sediments. it will be possible to develop and add new modules and functions including improved tools for importing and extracting data and maps. conclusions compared to the previous pda-based system, data entry in afieldwork app is much easier and faster with the new android system. the design is simpler and more intuitive. the presence of a real database system and look-up tables on the devices ensures data integrity. it also allows swift transfer of data to central databases after field work. the sqlite databases are binary and compatible between different platforms so an sqlite database generated on an android device can simply be moved to a windows computer and used from this without any conversion. the android system is easy to configure and the large open-source community makes a lot of free software available. geus has successfully used small digital handheld devices and software for digital data capture and observations on localities, rocks, samples, etc. in greenland for a number of years. with the development of the new afieldwork app, geus has modernised and improved this important field work tool for the geologist. acknowledgements the android system was tested during field work in south-east greenland by participants in the segment 2012. the participants in these field parties provided valuable input and comments to the system which are greatly appreciated. thanks to the two referees, thorsten nagel and denis m. schlatter, who contributed with valuable suggestions which greatly improved the manuscript. references schlatter, d.m., buller, g., larsen, u. & stensgaard, b.m. 2010: digital field data capture: the geological survey of denmark and greenland experiences in greenland. the association of applied geochemistry, explore 147, 2–14. asqlitemanger, rasterlite, spatialite http://www.gaia-gis.it/gaia-sins/ spatialite-sql-4.4.0.html mapsforge https://github.com/mapsforge/mapsforge authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen, denmark, e-mail mh@geus.dk geological survey of denmark and greenland bulletin 15, 2008, 53-56 53 the structure of the lithosphere under denmark has been investigated in relation to adjacent regions of sweden and germany. the most interesting result of the study is that the 120 km thick lithosphere under denmark appears to be a stretched version of the swedish lithosphere, which is more than twice as thick. during the international project tele seismic tomography across the tornquist zone (tor), field work and international interpretation were carried out between 1996 and 2002. following the field work period, modelvelocity computations were undertaken based on ob ser vations of distant earthquakes (e.g. arlitt 1999; shomali et al. 2002; voss et al. 2006), and recently an evaluation of the tor results was completed (nielsen 2007). the tor project investigates deeper parts of the earth than previous projects, and in particular the depth interval 50–300 km, which is below the crystalline crust. the investigations have included many geophysical features such as teleseismic p-wave tomography, rayleigh wave velocities, shear wave splitting and wave scattering. we have distinguished between relatively highand low-velocity zones, which also show variations in anisotropy and scatter characteristics. generalised high-velocity zones correspond to the lithosphere, while generalised relatively low-velocity zones are equivalent to the asthenosphere. the main outcome of the combined studies is that the deep lithosphere can be divided into three blocks separated approximately along the national boundaries between sweden and denmark and between denmark and germany. the boundaries between the blocks are steep, almost vertical. the denmark block has lithosphere properties between those to the north and south. based on previous crustal studies and the tor results, we suggest that the denmark block has evolved by stretching. the details in the new evaluation are derived from teleseismic tomography. here we present a synthesis of the many derived models in the light of the new evaluation. lower lithosphere: data and interpretation the area of investigation is shown in fig. 1 and examples of analyses of p-wave data are presented in fig. 2. the first and most important part of the data treatment is correction for the rather well-known structure of the crust, see fig. 2, middle frames (e.g. eugeno-s working group 1988; thybo 2000). the end products of the analyses are two patterns of travel-time residuals for the lower lithosphere (bottom frames of fig. 2). together with approximately 50 similar patterns of p-wave travel-time residuals, each for one earthquake, these constitute the data for the model-velocity computations. several tomographic velocity calculations have been made in 2-d cross-sections along the line of seismographs shown in fig. 1 (e.g. arlitt 1999; busche 2001; shomali et al. 2002; nielsen 2007). the results of nielsen’s (2007) work are shown in figs 3 and 4, which are, respectively, a cross-section with p-wave velocity deviations from a laterally homogeneous earth, and the horizontal gradients of these p-wave velocities. figure 4 pinpoints the sharpest boundaries be tween different lithosphere/asthenosphere char acte ris tics. the teleseismic tomography has a poor resolution in the upper 50 km of the earth, and the interval from 0 to 50 km evidence of stretching of the lithosphere under denmark søren gregersen, lene vandur nielsen and peter voss © geus, 2008. geological survey of denmark and greenland bulletin 15, 53–56. available at: www.geus.dk/publications/bull 58°n 4°e 12°e 20°e 54°n 50°n 200 km broadband seismographs short period seismographs denmark norway sweden poland germany c entral g raben elbe line ringkøbing–fyn high o slo g ra be n the netherlands sorgenfrei–tornquist zone fig. 1. the location of seismographs during field work of the tor project 1996–1997. the 2-d interpretation profiles of figs 3 and 4 follow the line in the middle of the cloud of seismographs from 59 °n, 16°e to 51°n, 9°e, crossing the sorgenfrei–tornquist zone at a right angle. the ringkø b ing–fyn high is a basement high separating sedimentary basins to its north and south. the elbe line is a geophysically recognised lineament. other important broad-scale features are the permian oslo graben and the mesozoic central graben. depth is not shown on figs 3 and 4. however, many previous investigations based on other data have provided rather detailed information about shallower structures (arlitt et al. 1999; pedersen et al. 1999). a dipping crustal transition from the baltica crust in the north-east to the avalonia crust in the south-west has been interpreted in the broad box a area. the blue and red p-velocity anomalies are computed with reference to a one-dimensional global travel-time model. two sharp and steep velocity changes, or red to blue steps, are seen in figs 3 and 4, and are outlined by boxes b and c. a third step at 57°n, seen as a colour change from light blue to dark blue, which was discussed by gregersen et al. (2006), may not be significant (fig. 4). the uncertainties of the exact locations of the transitions and their slopes are illustrated by the sizes of the boxes (b and c) in the transition regions. a low-velocity asthenospheric layer, as deduced from higher mode surface-wave studies mentioned by gregersen et al. (2006), is indicated by box d. this low-velocity layer does not extend to the north-east. this is consistent with the earlier interpretation of fundamental mode rayleigh waves that showed absence of low-velocity asthenospheric layers below the baltic shield at this depth (cotte et al. 2002). it can be difficult to evaluate which features in the models should be regarded as well defined and appropriate to interpret in geological terms. issues of spatial resolution, accuracy and uniqueness of the models are important. within the tor project considerable efforts have therefore been dedicated to examining these issues using, for example, tomographic inversions. in such an inversion it is straightforward to calculate the variance of model parameters, which quantifies the reliability of each parameter. this variance of a model parameter has only limited value, partly because mathematical assumptions of linearity are simplified, and partly because the often complicated interactions between model parameters are ignored. the estimation of resolution also depends on choices in the mathematical inversion procedure. the resolution is limited by station spacing, ray geometry between earthquakes and seismic stations, and the frequency content of data, and it is different for various parts of the model. these complex problems mean that different approaches must be considered to quantify reliability. this has been carried out in previous analyses and evaluated by nielsen (2007). suggested evolution of the lithosphere under denmark the broad-scale division of the lithosphere into three different blocks, with denmark as the middle one of intermediate structure, has been confirmed by several studies. one major outcome of the tor project is therefore unquestionably that the sorgenfrei–tornquist zone is connected to a very deep and prominent lithospheric velocity difference (box c in figs 3 and 4). another marked change in lithosphere-asthenosphere properties is found in the depth range 0–120 km between the southern part of the ringkøbing–fyn high and 54 400 200 0 y (k m ) –200 –400 c al cu la te d cr us ta l a nd up pe r m an tle e ffe ct s o bs er ve d tr av el tim e re si du al s li th os ph er ic a nd as th en os ph er ic e ffe ct s 400 200 0 y (k m ) –200 –400 400 200 0 y (k m ) x (km) –200 –400 –200 0 200 x (km) –200 0 200 re si du al s (s ) 1.5 1 0.5 0 –0.5 –1 –1.5 –1.5 re si du al s (s ) 1.5 1 0.5 0 –0.5 –1 –1.5 re si du al s (s ) 1.5 1 0.5 0 –0.5 –1 fig. 2. p-wave travel time residuals from two earthquakes, in japan (19 october 1996, 32°n 132°e, wave arrival to tor area from north-east) and in mexico (11 january 1997, 18°n 103°w, wave arrival to tor area from north-west), each in one column. for each of the earthquakes the upper diagram shows total observed residuals (observed arrival times minus expected arrival times according to the global average tables iasp91). the middle diagram shows computed crustal residuals to 50 km depth (from pedersen et al. 1999), which, when subtracted from the observed ones, yield the lower lithosphere residuals below 50 km depth which are shown in the lower diagram (from pedersen et al. 1999). the elbe line (boxes a and b in figs 3 and 4). these interpretations together with the well-known surface geology and crustal structure (e.g. eugeno-s working group 1988; thybo 2000) have been discussed at many europrobe meetings. the outcome of these discussions is that the evolution of the lithosphere-asthenosphere system can be summarised in graphical form (fig. 5), with collision, spreading, shearing and compression over the last 500 million years. based on the tor-project studies by cotte et al. (2002) and gregersen et al. (2002, 2006) we suggest that together with the other geophysical data the surface wave results and the tomography results (figs 3 and 4) indicate a lithosphere thickness of a little less than 100 km in the south-western part of the profile, a little over 100 km in the middle, and a lithosphere thickness of more than 200 km in the north-eastern part of the region. the baltic shield in sweden, i.e. the part of the craton exposing old, precambrian crystalline rocks, terminates to the south-west at the sorgenfrei–tornquist zone. from deep dril ling data and seismic crustal studies, it is deduced that material originating from the craton extends southwards from the sorgenfrei–tornquist zone, gradually thins where overlying sediments are thick and terminates just south of the ring købing–fyn high. in the late precambrian to early palae ozoic (c. 600 ma), the eastern european craton including the baltic shield formed the core of the palaeo-continent baltica, and the tor study region was a passive continental margin. the tomographic images reveal the structure originating from this episode, reworked by a number of later events, as illustrated by fig. 5. the evolution can be divided into four stages. during stage 1 two separate lithospheric plates collided in late palaeozoic times. in stage 2 the area was deformed by lithospheric stretching and transcurrent faulting, the ringkø bing–fyn high with its thick crust was detached from the baltic craton, and the various blocks of the ringkø bing–fyn high rotated slightly. in the ensuing stage 3, lithospheric stretch ing was perpendicular to the trend of the plate transition and the ringkø bing–fyn high. the compressional stage 4 is very different, involving inversion of the sorgenfrei–tornquist zone. conclusions our essential contributions from geophysics to the geological evolutionary account are the measurements of crustal and lithospheric thinning. in previous studies of the crustal structure (eugeno-s working group 1988) it was suggested that the crust in denmark between the sorgenfrei–tornquist zone and the ringkøbing–fyn high had been stretched during the late palaeozoic and the mesozoic. preliminary modelling of the stretching of the sedimentary basins and the crystalline crust has been carried out by nielsen & balling (1990). fuller modelling, including the entire lithosphere, has been undertaken by frederiksen et al. (2001). since the results of the tor project also show a distinct thinning of the lower lithosphere from the swedish to the danish area, we see this as evidence of essentially full lithospheric stretching. the reasoning has developed historically from recognition of stretching in sedimentary basin evolution, and then stretching of the crust (eugeno-s working group 1988), to the present claim of stretching of the entire lithosphere. this account can now be reversed in a well-founded story of cause and effect: the lithosphere is stretched, and the crystalline crust is thereby stretched, creating the sedimentary basins. 55 100 200 d ep th ( km ) 300 51 52 53 –8 –6 –4 –2 0 2 4 6 8 54 55 latitude (deg) p velocity perturbation (%) 56 57 58 c b a d el rfh stz 100 200 d ep th ( km ) 300 51 52 53 0.000 0.005 0.010 54 55 latitude (deg) absolute horizontal p velocity gradient (1/s) 56 57 58 c b a d el rfh stz fig. 3. tomographic image of the lithosphere/asthenosphere system in a profile along the tor array from 59°n, 16°e to 51°n, 9°e. blue colours are areas with high p-wave velocities, while red colours are areas with low pwave velocities with respect to the laterally homogeneous iasp91 model (kennett & engdahl 1991). the boxes show the uncertainty around the lateral changes (see text). stz, sorgenfrei–tornquist zone; rfh, ringkø bing–fyn high; el, elbe line. fig. 4. tomographic image along the tor array. for location see fig. 1. dark areas show the largest horizontal gradients in p-wave velocities, corresponding to large changes in material. the boxes show the uncertainty around the lateral changes (see text). references arlitt, r. 1999: teleseismic body wave tomography across the transeuropean suture zone between sweden and denmark, 109 pp. unpublished ph.d. thesis, swiss federal institute of technology, zurich, switzerland. arlitt, r., kissling, e., ansorge, j. & tor working group 1999: 3-d crustal structure beneath the tor array and effects on teleseismic wavefront. tectonophysics 314, 309–319. busche, h. 2001: manteltomographie: verspätungen von wellenzügen im norddeutsch–dänischen becken, 105 pp. unpublished ph.d. thesis, christian-albrechts-universität kiel, germany. cotte, n., pedersen, h.a. & tor working group 2002: sharp contrast in lithospheric structure across the sorgenfrei–tornquist zone as inferred by rayleigh wave analysis of tor1 project data. tectonophysics 360, 75–88. eugeno-s working group 1988: crustal structure and tectonic evolution of the transition between the baltic shield and the north german caledonides (the eugeno-s project). tectonophysics 150, 253–348. frederiksen, s., nielsen, s.b. & balling, n. 2001: a numerical model for the norwegian–danish basin. tectonophysics 343, 165–183. gregersen, s., voss, p., shomali, z.h. & tor working group 2002: summary of project tor: delineation of a stepwise, sharp, deep lithosphere transition across germany–denmark–sweden. tectonophysics 360, 61–73. gregersen, s., voss, p., shomali, z.h., grad, m., roberts, r.g. & tor working group 2006: physical differences in the deep lithosphere of northern and central europe. in: gee, d.g. & stephensson, r.a. (eds): european lithosphere dynamics. geological society memoirs (london) 32, 313–322. kennett, b.l.n. & engdahl, e.r. 1991: traveltimes for global earthquake location and phase identification. geophysical journal interna tional 105, 429–465. nielsen, l.v. 2007: teleseismic p wave tomography across the sorgen frei–tornquist zone. evaluation of tor project results, 109 pp. un published cand. scient. thesis, university of copenhagen, denmark. nielsen, s.b. & balling, n. 1990: modelling subsidence, heat flow, and hydrocarbon generation in extensional basins. first break 8, 23–31. pedersen, t., gregersen, s. & tor working group 1999: project tor: deep lithospheric variation across the sorgenfrei–tornquist zone, southern scandinavia. bulletin of the geological society of denmark 46, 13–24. shomali, z.h., roberts, r.g. & tor working group 2002: non-linear body wave teleseismic tomography along the tor array. geophysical journal international 148, 562–574. thybo, h. 2000: crustal structure and tectonic evolution of the tornquist fan region as revealed by geophysical methods. bulletin of the geo logical society of denmark 46, 145–160. voss, p., mosegaard, k., gregersen s. & tor working group 2006: the tornquist zone, a north-east-inclining lithospheric transition at the south-western margin of the baltic shield: revealed through a non-linear teleseismic tomographic inversion. tectonophysics 416, 151–166. 56 authors’ addresses s.g. & p.v., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: sg@geus.dk l.v.n., niels bohr institute, university of copenhagen, juliane maries vej 30, dk-2100 copenhagen ø, denmark. n s r fh 1. ordovician–devonian: coincident with caledonian orogeny 2. late carboniferous – early permian: coincident with variscan orogeny 3. triassic–jurassic: lithospheric stretching 4. late cretaceous – early cenozoic: coincident with alpine orogeny r fh r fh st z fig. 5. generalised summary diagram of the large-scale geological development of the tor area (fig. 1). blue shows areas with sediments, orange shows areas of crystalline crust, and red is the uppermost mantle lithosphere. the sorgenfrei–tornquist zone (stz) and the ringkøbing–fyn high (rfh) act through time as compression, spreading and shearing zones. the arrows show the regional stress field (from gregersen et al. 2006). geological survey of denmark and greenland bulletin 12, 63-73 63 the lower part of the overlying main body of greenish and greyish mudstones in the offshore succession can be correlated with the coeval and lithologically similar upper part of the lillebælt clay formation. the upper part of the horda formation can be correlated with the søvind marl formation, which consists of grey marls. the highest part of the horda formation, only observed in central graben wells, may be correlated with the viborg formation on biostratigraphic evidence. hefring member new member history. the hefring member consists of sandstone deposits within the horda formation. these sandstones have not previously been recognised as a separate unit in the danish sector. derivation of name. after the goddess hefring. type well. danish sector well floki-1, 1793.4–1731.3 m mdrt (fig. 53). distribution and thickness. the hefring member is only known from the floki-1 well located in the northern part of the danish sector. as the unit currently cannot be identified on seismic sections, its further distribution is unknown. in the floki-1 well, the member is 62 m thick. lithology. the hefring member consists of greenish grey, fine-grained, immature sandstones with glaucony grains. logcharacteristics.thehefring member ischaracterisedby aconspicuousblockysignatureonthegamma-ray, sonicand density logs (fig. 53). gamma-ray responses are lower than those of the enveloping horda formation mudstones. the hefring member can also be recognised from a combination of the density and neutron logs as the presence of pure sandstones results ina ‘cross-over’of the two logcurves (fig. 53). boundaries. the boundaries with the mudstones of the horda formation are sharp and characterised by prominent shifts on the gamma-ray and sonic logs (fig. 53). depositional environment. no cores have been taken in the hefring member, but the sandstones were probably deposited from concentrated gravity flows, based on log similarity with the other fine-grained sandstone bodies in the nearby siri canyon. age. lutetian (middle eocene) based on the age of the associated horda formation mudstones. correlation. based on biostratigraphic data, the hefring member may be contemporaneous in part with the lillebælt clay formation onshore denmark, with the lower part of the grid sandstone member (knox & holloway 1992) in the viking graben and with the upper part of the tay sandstone member (knox & holloway 1992) in the northern part of the central graben. westray group the westray group is the upper of the two groups established by knox & holloway (1992) to replace the hordaland group of deegan & scull (1977; fig. 3). in the central north sea and in the danish sector of the north sea, the westray group is represented by the lark formation. lark formation history. the lark formation was established by knox & holloway (1992) for the brownish grey mudstone-dominated lithofacies of the westray group that overlies the more variable association of red and green-grey mudstones, silty mudstones and sandstones of the horda formation and underlies the grey, sandy and shelly mudstones, siltstones and sandstones of the nordland group of deegan & scull (1977; fig. 3). the lark formation is also recognised in the danish sector although its lithology is more variable than that given in the original description. type well. british sector well 21/10-4, 1867–1217 m mdkb. danish reference wells. mona-1, 2363.5–1598.3 m mdkb (fig. 46); siri-1, 1916.5–819.3 m mdkb (fig. 54). distribution and thickness. the lark formation extends over the central and northern north sea and is probably present in the entire danish sector of the north sea. its depocentre is in the central and northern part of the danish sector, along the eastern boundary of the danish central graben, where it reaches a thickness of 1194 m in the siri-3 well. the lark formation thins west to a thickness of 389 m in the tordenskjold-1 well in the central graben, and east to a thickness of 240 m in the s-1 well on the ringkøbing–fyn high (fig. 55). 64 lithology. the lower lark formation (l1–3, see below) is dominated by dark, greenish grey, non-fissile mudstones in most wells; in some wells subordinate intervals of brownish grey mudstones are also present. thin layers of white or reddish brown carbonate are also recorded in the upper levels of the lower lark formation. the upper lark formation (l4, see below) is dominated by pale to dark brownish grey mudstones with subordinate intervals of greenish grey mudstones in its lower levels. the uppermost 50–100 m of the formation consist of yellowish grey to light brown mudstones. in eastern and northern parts of the danish sector, discrete sandstoneinterbeds and thin sandstone stringers occur throughout the formation. log characteristics. the lower part of the lark formation is characterised by an overall stable gamma-ray log signature, whereas the upper part of the formation has a more unstable signature (figs 46, 54). this change in gammaray log signature coincides approximately with the change from lithologies dominated by greenish grey mudstones to lithologies dominated by dark to light brownish grey mudstones at the base of unit l4 (see below). boundaries. the base of the lark formation is marked by a change from fissile, greenish grey mudstones of the horda formation to non-fissile, greenish grey mudstones of the lark formation. this change in lithology coincides with an abrupt increase in gamma-ray values to a consistently higher level than that displayed by the horda formation (figs 46, 51, 54). wells in the eastern part of the danish central graben and on the ringkøbing–fyn high show a conspicuous log break on the gamma-ray log at the formation boundary, whereas the log break is less pronounced in wells from the central and western parts of the danish central graben (fig. 51). although the actual increase in gamma-ray response may be limited in the latter wells, the offset is usually sharp and well defined. on the sonic log, the boundary between the horda formation and the lark formation is characterised by a transition from a stable sonic signature to one characterised by numerous fluctuations. the lark formation is overlain by the undifferentiated nordland group of deegan & scull (1977). over most of the area, the boundary seems conformable and is represented by a change from yellowish grey and light brown mudstones to medium to dark grey mudstones characte1916.5 819.3 horda fm lark fm nordland gp l4 l3 l2 l1 balder fm sele fm lista fm våle fm chalk gp 800 900 1000 1100 1200 1300 1400 1500 1600 1700 1800 1900 2000 2100 2200 m siri-1 gr sonic no log data fig. 54. siri-1, danish reference well for the lark formation. the units l1–4 are all present in this well. black bars show cored sections. 65 rised by intervals with shell-hash and coarse-grained sands. this boundary is marked by a conspicuous gamma-ray peak at the base of a 20–40 m thick interval with elevated gamma-ray values in the lowermost nordland group (figs 46, 51). this interval is further characterised by a marked double peak on the gamma-ray log. in thenorth-easternparts of the danish sector (nini-1, vanessa-1, cecilie-1 and siri-1; figs 1, 51, 54), sediments of the nordland group rest unconformably on the lark formation. in this area the uppermost lark formation and the lowermost nordland group are missing, probably due to erosion and/or non-deposition. the distinct gamma-ray peak that marks the top of the lark formation as well as the double gamma-ray peak in the lowermost nordland group are lacking in these wells, and therefore the top of the lark formation is more difficult to identify on petrophysical logs. subdivision. the lark formation can be subdivided into four major mudstone packages, l1–l4, based on seismic and log evidence (figs 46, 49, 51, 54, 56; plates 1–5). these units are described below; isochore maps of the units are shown in fig. 57a–d. l1 (figs 51, 56a, b, 57a; plates 1, 4, 5) this unit has been recognised in the north-eastern part of the danish sector only (fig. 57a). it is bounded beneath by the th marker and above by the tl1 marker (fig. 56a, b). it is characterised by downlapping reflectors and represents a south-westwards prograding mudstone succession. on the gamma-ray log, the l1 unit is characterised by a relatively high and relatively stable response. in most wells, it shows a weakly concave pattern, going from a relatively high gamma-ray response at its base, over a gamma-ray low halfway through the unit to a level close to starting level at the top of the unit (e.g. ida-1, inez-1, k-1, f-1 and sandra-1; plates 1, 4, 5). in the siri-1 and siri-3 wells, near the south-western limit of the l1 unit, the gamma-ray log motif instead appears slightly convex (figs 51, 54; plates 1, 4). the l1 unit consists predominantly of greenish grey mudstones but also includes yellowish brown and dark grey mudstones. l2 (figs 46, 51, 56a, b, 57b; plates 1–5) the unit is recognised over the entire study area. on the gamma-ray and sonic logs the unit is characterised by a stable log signature. the gamma-ray log shows two to three slightly concave patterns with signatures similar to that of the l1 unit (fig. 51; plates 1, 4). the lithology is characterised by dark beige-grey to greenish grey mudstones, greenish colours becoming dominant towards the top of the unit. l3 (figs 46, 51, 56a, b, 57c; plates 1–5) this unit is encountered in the northern and eastern parts of the danish sector, east of the central graben (fig. 57c) but is not recognised on logs or seismic sections in the central graben area. it is characterised by stable gamma-ray and sonic log signatures (figs 46, 51; plates 1–4). the unit consists almost invariably of dark, greenish grey mudstones. l4 (figs 46, 51, 56a, b, 57d; plates 1–5) the unit is recognised over the entire study area (fig. 57d). the interval is characterised by a slightly more unstable gamf-1 inez-1 siri-1 francisca-1 frida-1 tordenskjold-1 s-1 mona-1 25 km 200 300 400 500 600 700 800 900 1000 1100 1200 1300 freja member distribution dufa member distribution lark formation thickness (m) fig. 55. isochore map of the lark formation in the study area. the positions of the two danish reference wells, mona-1 and siri-1, are indicated in the figure. the map also shows the distribution of the sandstones of the dufa and freja members and the location of their type and reference wells, inez-1 and f-1, and francisca-1 and frida-1, respectively. 66 fig. 56. a: sw–ne-trending seismic section (rtd81-re94-22a) showing the complex architecture of the lark formation and its subdivision into l1–4 units. the vertical white line indicates change in section direction. the locations of the two seismic sections are shown on fig. 1; abbreviations as in figs 49 and 51. 0 500 1000 1500 2000 2500 3000 3500 0 500 1000 1500 2000 2500 3000 3500 tordenskjold-1 f-1 tabita-1 cecilie-1 siri-3 sw ne horda fmhorda fm l2l2 lark fmlark fm l4 horda fm l2 aa l2l2 l1 lark fmlark fm l4 l3 l1 10 km lark fm lark fm l4 l3l3l3 l4 l3 uou tl2 th tb tc tl tf uou tl2 tl uou tl2 tl1 th tb tc twt (msec) twt (msec) 67 25 50 75 100 125 thickness (m) l1 25 km 25 km 100 200 300 400 thickness (m) l2 100 200 300 400 500 600 700 800 thickness (m) 100 200 300 400 500 600 700 800 thickness (m) l3 l4 a b c d fig. 57. isochore maps of lark formation subunits. a: l1. b: l2. c: l3. d: l4. a and d are at the same scale, b and c are at the same scale. 1500 2000 2500 3000 3500 twt (msec) 20 km nini-2 d-1 nolde-1 nw se tl uou tl2 tl1 tb tc l1 l2l2 l3 l4 l1 l2 l3 l4 bb fig. 56. b: nw–se-trending seismic section (rtd81-re94-14a) showing subdivision of the lark formation and marked thinning of this formation towards the south-east. the locations of the two seismic sections are shown on fig. 1; abbreviations as in figs 49 and 51. 68 ma-ray and sonic log signature than that of the underlying units (figs 46, 51; plates 1–5). it is dominated by brown to yellowish brown mudstones, but in some wells an interval of greenish grey mudstones occurs in its lower part. in wells to the east and north, thin sandstones are interbedded with the mudstones and become more frequent towards the top of the unit. two thick sandstone units occur in the lark formation on the ringkøbing–fyn high and are described here as two new members (dufa and freja members). macroand ichnofossils. only observed in cores taken in the freja member (see below). microfossils and palynomorphs. farthest to the north and east the basal part of the formation includes the downhole succession of ho areosphaeridium diktyoplokum and ho a. michoudii indicating a late priabonian (late eocene) age for the base of the formation in this area. in the central graben area, the base of the lark formation is significantly younger. here it contains an event succession characteristic of the middle and lower rupelian (lower oligocene) stage (hos of phthanoperidinium amoenum, achilleodinium biformoides and phthanoperidinium comatum). the top of the lark formation is bracketed by a number of conspicuous biostratigraphic events: the uppermost part contains the hos of the benthic foraminifers asterigerina staeschei and elphidium inflatum followed downhole by the ho of uvigerina tenuipustulata. dinoflagellate events near the top of the lark formation include the hos of apteodinium spiridoides and cousteaudinium aubryae. the lowermost part of the overlying nordland group contains the hos of the calcareous microfossils bolboforma clodiusi, bolboforma spiralis and bolboforma metzmacheri, the ho of the benthic foraminifer bulimina elongate, and the ho of the dinoflagellate cyst cannosphaeropsis passio. a large number of hos characterise the lark formation; key events are listed in fig. 5c. depositional environment. the l1 unit is characterised by abundant agglutinated foraminifers dominated by rhabdammina discreta and similar tubular taxa, together with haplophragmoides spp. and recurvoides spp. the microfaunal assemblage indicates that the unit was predominantly deposited in an open marine, dysoxic palaeoenvironment at upper bathyal depths. the l2 unit and most of the l3 unit are characterised by an increasing abundance and diversity of calcareous benthic and planktonic foraminifers. the relative proportions of agglutinated, calcareous planktonic and benthic foraminifers vary considerably from well to well, indicating pronounced lateral changes in the depositional environment. the calcareous plankton/benthos ratio is usually low, indicating a neritic setting for most of the succession, but in a few restricted intervals it may reach 1:2 or even 1:1 indicating an outer neritic setting. thus, the foraminifer assemblage indicates an open marine, neritic to outer neritic setting with well-oxygenated bottom conditions for the lower to middle part of the lark formation. the microfossil assemblage in the uppermost part of the l3 unit as well as the l4 unit is dominated by calcareous benthic foraminifers, and agglutinated foraminifers are generally rare. epifaunal and shallow infaunal foraminifers are more common than deep infaunal taxa, indicating oxic bottom conditions during this interval. in general, the microfaunal assemblage in this part of the lark formation suggests that it was deposited in a neritic, probably middle neritic, palaeoenvironment over most of the study area. the palynofacies assemblage in the lark formation is characterised by a rich dinoflagellate assemblage and abundant dispersed terrestrial matter (phytoclasts, spores and pollen), indicating an open marine environment with considerable influx from nearby land areas. stratigraphic variations in the relative abundance of terrestrial palynomorphs in the lark formation suggest successive pulses of progradation and backstepping of the palaeocoastline. age. the lark formation is of priabonian to serravallian (late eocene to middle miocene) age with eocene sediments being present in the l1 unit only. the l1 unit is priabonian to early rupelian (early oligocene) in age, the base of the unit being oldest farthest to the north and east and younging towards the south and west. the age of the l2 unit is rupelian; the l3 unit is rupelian in age in its lower part and chattian (late oligocene) in its upper part. the rupelian–chattian boundary is located in the lower part of the unit. the chattian–aquitanian (oligocene– miocene) boundary is located just above the top of the l3 unit. in some wells, a hiatus is indicated at this level by the clustering of hos. the chattian–aquitanian, aquitanian–burdigalian, burdigalian–langhian and langhian–serravallian stage boundaries are all located in the l4 unit. the uppermost part of the lark formation is of midserravallian age. correlation. based on biostratigraphic correlation, the lowermost l1 unit is probably largely coeval with the viborg formation onshore denmark, and with sequence 4.1 of michelsen et al. (1998). the l2 unit may be corre69 lated with the linde clay onshore denmark (informal mudstone unit described by heilmann-clausen 1995). intervals in the l3 unit may be correlated with the branden clay (ravn 1906) onshore denmark, based on lithological similarities and biostratigraphy. intervals in the lark formation around the l3–l4 boundary (around the chattian–aquitanian boundary) may be correlated with the two lowermost, clay-rich units of the onshore vejle fjord formation(thebrejningclayandvejlefjordclayoflarsen & dinesen1959).theuppermostpart of the lark formation possibly correlates with the onshore arnum formation (sorgenfrei 1958) and the hodde formation onshore denmark (rasmussen 1961), based on biostratigraphy. dufa member new member history. the dufa member comprises a thick sandstonedominated unit that occurs within unit l3 of the lark formation in the northern and eastern part of the danish sector of the north sea. the unit has not been previously described. derivation of name. after the goddess dufa. type well. danish sector well inez-1, 697.1–485.5 m mdkb (figs 41, 58; plates 1, 5). reference well. danish sector well f-1, 337.5–324.3 m mdkb (figs 58, 59; plate 5). distribution and thickness. the dufa member is present in the north-eastern part of the danish sector of the north sea (fig. 55). in its type well, the dufa member is 210 m thick and consists of three major sandstone units with thicknesses 30–120 m (fig. 41). the sandstone units are separated by mudstone intervals up to 20 m thick. towards the north, in the f-1 well, the lower sandstone units are fig. 58. n–s-trending composite seismic section (rtd81-re94-45/rtd81-re94-09). blue-coloured lines indicate the outline of the dufa member. the gamma-ray logs from the inez-1 and f-1 wells are inserted (see figs 55 and 59 for depth-converted gamma-ray logs for the two wells). the horda formation is thin in this area and the top horda reflector (th) is therefore indistinguishable from the top balder reflector (tb). the location of the seismic section is shown in fig. 1; abbreviations as in fig. 49. inez-1f-1 0 500 1000 twt (msec) tf tl2 tb/th n s dufa mb dufa mbdufa mb 5 km 70 well, the lower sandstone intervals of the dufa member show the presence of a number of 5–10 m thick sandstone packets showing blocky, decreasing-upwards gamma-ray log signatures suggesting coarsening-upwards sand bodies. these sandstones are separated by intervals of fig. 60. francisca-1, type well for the freja member. black bars show cored sections. 1400 1500 1600 1700 1800 m 1562.8 1840.7 freja mb lark fm francisca-1 gr sonic missing (figs 58, 59). the dufa member is absent in wells west of f-1. lithology. the lower sandstone units predominantly consist of coarsening-upwards successions of very fine-grained to fine-grained, greenish brown, muddy sandstones. the upper sandstone unit fines upwards and consists of mediumto coarse-grained, quartzitic, relatively pure sand with intervals rich in glaucony. lignite has been observed in cuttings samples. log characteristics. the member is characterised by an overall blocky signature on the gamma-ray log. in the type fig. 59. f-1, reference well for the dufa member. dufa mb lark fm 300 400 500 600 m f-1 324.3 337.5 gr sonic 71 mudstones with higher gamma-ray response. the upper unit is characterised by an overall blocky signature with minor gamma-ray peaks and trends suggesting a number of fining-upwards intervals, 10–20 m thick, and a few coarsening-upwards intervals, 5–10 m thick (fig. 41). boundaries. the upper and lower boundaries of the member with the mudstones of the lark formation are sharp and characterised by prominent shifts on the gamma-ray log (figs 41, 59). depositional environment. judging from seismic evidence, the dufa member is positioned partly on the offlap break, partly seaward of it (fig. 58). based on this palaeosetting and the presence of lignite in cuttings samples, the dufa member sandstones are interpreted to represent deltaic, shallow-marine sediments, probably deposited in pulses during an overall relative sea-level low. age. rupelian, based on the age of the enveloping mudstones. correlation. there are no danish onshore correlatives to the dufa member. the correlation with the norwegian offshore successions is currently uncertain. freja member new member history. the freja member is a conspicuous sandstonedominated unit that occurs within the upper levels (l4) of the lark formation in the northern and central parts of the danish sector of the north sea. the unit has not previously been described. derivation of name. after the goddess freja. type well. danish sector well francisca-1, 1840.7–1562.8 m mdrt (figs 60, 61). reference well. danish sector well frida-1, 1623.5–1487.7 m mdrt (fig. 62; plate 4). distribution and thickness. the freja member is present in the northern and central parts of the danish sector of the north sea (fig. 55). in its type well, the freja member spans a stratigraphic interval of c. 280 m and includes major sandstone units separated by subordinate intervals of mudstones (fig. 60). in the cecilie-1 well, the member is c. 150 m thick whereas in the frida-1 well the member attains c. 130 m (fig. 62). lithology. in its type well, the lower half of the freja member consists of very fine-grained to fine-grained quarzitic sandstones with many thin mudstone interbeds. the member becomes less muddy in the upper third of this interval. the upper third of the freja member consists largely of relatively pure quarzitic, very fine-grained sandstones with mudstone interbeds becoming frequent towards the top (fig. 60). between these two major sandstone units is a c. 60 m thick interval dominated by mudstones but with carbonate-cemented, sandstone-dominated packets in its upper part. log characteristics. the freja member has an overall blocky gamma-ray log signature. in the type well, its lowermost part (1840.7–1750 m) can be split into a number of smaller units with blocky or increasing-upwards gamma-ray log signatures separated by gamma-ray peaks. in comparison, the overlying sandstones (1750–1720 m) display a more stable, low gamma-ray log pattern with few gamma-ray log spikes (fig. 60). the mudstone-dominated interval (1720–1660 m) separating the two sandstone-dominated units in the type well generally shows high gamma-ray values: thin calcite-cemented sandstone packets are intercalated with the mudstones in this interval (e.g. 1680– 1670m)and show decreasing-upwards gamma-ray values. boundaries. the lower boundary of the freja member with the lark formation mudstones is sharp and characterised by prominent shifts on the gamma-ray and sonic logs. in the type well, where the upper levels of the freja member are characterised by interbedded mudstones and sandstones, the upper boundary of the member is less prominent. in this well, it is placed at the top of the uppermost discrete sandstone bed, at 1562.8 m (fig. 60). macroand ichnofossils. intervals with shell debris have been observed in core sections of the freja member in the francisca-1 well. ichnofossil genera from the freja member comprise chondrites ispp., phycosiphon ispp., planolites ispp., terebellina ispp., thalassinoides ispp. and zoophycos ispp. depositional environment. the freja member represents stacked successions of thickand thin-bedded turbidite sands deposited in submarine channels and proximal levee environments (figs 60, 62). the upper parts of the turbidite successions show transitions from normally graded turbidites, deposited in slightly more distal levee environ72 francisca-1 w e n s 2 km 0 500 1000 1500 2000 2500 3000 twt (msec) freja memberfreja memberfreja member horda fmhorda fm lark fmlark fm lark fmlark fm horda fm lark fm l3l3 l2l2 l4l4 lark fm tl tf uou tl2 th tb ments and minor turbidite channels, to mainly silty turbidite deposits that represent distal levee and fan fringe environments and the transition to the open slope. the source of the sand was probably a marginal marine shelf environment, judging by the abundance of the marginal marine acritarch paralecaniella indentata. age. in the type well, the freja member is chattian to aquitanian in age, based on the age of mudstones within and bounding the member. in the frida-1 well, the freja member is entirely chattian in age. correlation. the freja member is broadly contemporaneous with the vejle fjord formation onshore denmark, with the vade formation (hardt et al. 1989) in the norwegian central graben and with the skade formation (hardt et al. 1989) in the viking graben. thick, coarsening-upward sandstone bodies are present above the dufa memfig. 61. composite seismic section (dk1–5623a re94/dk1–0448b re94) with the freja member indicated between the reflectors uou and tf. the gamma-ray log from the francisca-1 well is indicated on the figure (see fig. 60 for depth-converted gamma-ray log). the vertical white bar indicates where the section changes direction. the location of the seismic section is shown on fig. 1; abbreviations as in fig. 49. 73 1487.7 1623.5 freja mb lark fm 1400 1500 1600 1700 m frida-1 gr sonic ber in the inez-1 well (shown as unnamed sandstones in fig. 2); these sandstones may be contemporaneous or even contiguous with those of the freja member. however, confident correlation on the basis of log and seismic data is not possible at present. acknowledgements aage bach sørensen (geus) is thanked for help with seismic interpretations. yvonne desezar, johnny e. hansen and birthe amdrup are thanked for preparation of microfossil and palynology samples. the referees robert w.o’b. knox (british geological survey) and paul van veen (conocophilips norway) are thanked for their constructive criticism of the manuscript; the editorial contributions of adam a. garde, jon r. ineson and martin sønderholm are gratefully acknowledged. this work was made possible through grants from the danish energy authority, under the energy research project framework 2000. fig. 62. frida-1, reference well for the freja member. black bar shows cored section. references ahmadi, z.m., sawyers, m., kenyon-roberts, s., stanworth, c.w., kugler, k.a., kristensen, j. & fugelli, e.m.g. 2003: paleocene. in: evans, d. et al. (eds): the millenium atlas: petroleum geology of the central and northern north sea, 235–259. london: geological society. armentrout, j.m., malecek, s.j., fearn, l.b., sheppard, c.e., nayler, p.h., miles, a.w., desmarais, r.j & dunay, r.e. 1993: log-motif analysis of paleogene depositional systems tracts, central and northern north sea: defined by sequence stratigraphic analysis. in: parker, j.r. (ed.): petroleum geology of northwest europe: proceedings of the 4th conference, 45–57. london: geological society. aubry, m.-p. et al. 2002: proposal: global standard stratotype-section and point (gssp) at the dababiya section (egypt) for the base of the eocene series, 58 pp. unpublished report, international subcommision on paleogene stratigraphy. aubry, m.-p. et al. 2003: chronostratigraphic terminology at the paleocene/eocene boundary. geological society of america, special paper 369, 551–566. berggren, w.a. & aubert, j. 1975: paleocene benthonic foraminiferal biostratigraphy, paleobiogeography and paleoecology of atlantic– tethyan regions; midway-type fauna. palaeogeography, palaeoclimatology, palaeoecology 18, 73–192. geological survey of denmark and greenland bulletin 15, 2008, 29-32 in denmark most of the water used in private households, in the industry and for irrigation in agriculture comes from subsurface aquifers. some of the most important aquifers in jyl land, western denmark, are sand layers deposited from 23 to 15 ma ago, in the early neogene (early to middle miocene). about 23 ma ago, in the early miocene, the coastline ran nw–se across present-day jylland (rasmussen 2004). global climatic variations led to major sea-level changes (zachos et al. 2001), which in combination with increased sediment transport from the north (the present norway) resulted in deposition of several huge, fluvio-deltaic sand systems intercalated with marine clay (e.g. rasmussen 1961; rasmussen 2004; rasmussen & dybkjær 2005). the geological survey of denmark and greenland (geus) and the regional environment centres (the former counties (amter)) in jylland are working in close cooperation to study the early neogene succession; the main purposes are: (1) to find new aquifers, (2) to map the extent of known aquifers and clarify their mutual relationships, in order to evaluate the size of the water resources and optimise production, and (3) to protect the aquifers from pollution due to leaching from the surface. in order to map the complex sedimentary succession, it has been necessary to combine several geological disciplines, including seismic interpretation, sedimentology, correlation of geophysical logs, and biostratigraphy (e.g. dybkjær 2004; ras mussen 2004; rasmussen et al. 2004; piasecki 2005; rasmussen & dybkjær 2005; dybkjær & rasmussen 2007). this article shows some results of a detailed dinoflagellate cyst stratigraphy, which is based on an extensive database (fig. 1). we present here for the first time a dinoflagellate cyst zonation for the complete neogene succession in the danish area. dinoflagellates and their cysts dinoflagellates are eukaryotic, single-celled organisms that occur as motile cells with two flagella, one which encircles the cell, and one longitudinal flagellum. these organisms, which also occur abundantly at present, include autotrophs, phagotrophs, symbionts and parasites. photosynthetic species (autotrophs) ac count for about half of the living genera and play an important role in the marine ecosystem as primary producers. some marine species produce toxins or cause red tides and several types of shellfish poisoning. dinoflagellates have left a rich fossil record, mainly of organic-walled cysts, in mesozoic and ceno zoic rocks; examples of fossil dinoflagellate cysts are shown in fig. 2. some dinoflagellate species form cysts as part of their life-cycle; others form cysts as a survival strategy, during unfavourable environmental conditions (e.g. low sea surface temperatures or lack of nutrients). the morphology of a cyst reflects the morphology of the motile dinoflagellate. the rapid evolution of the relatively complex, fossilisable dinoflagellate cyst wall makes these fossils ideal for biostratigraphic purposes (e.g. fensome et al. 1996). material and methods the dinoflagellate cyst zonation scheme presented here is based on data from more than 50 onshore and offshore boreholes and a series of onshore exposures (fig. 1). the sediment samples from the boreholes and exposures were processed at the palynological laboratory at geus using standard palynological preparation methods. a new neogene biostratigraphy for denmark karen dybkjær and stefan piasecki © geus, 2008. geological survey of denmark and greenland bulletin 15, 29–32. available at: www.geus.dk/publications/bull 29 9°e wells with biostratigraphy (49) wells without biostratigraphy (5) exposures (25) 50 km 53°n 55°n 57°n 8°e 10°e 12°e north sea jylland sweden germany denmark frida-1 lone-1 tove-1 s-1 100 km fig. 1. map of denmark showing the location of studied wells and exposures. the insert map shows the location of the offshore exploration wells included in this study. the danish neogene stratigraphy has traditionally been based on molluscs and foraminifers whereas early studies of dinoflagellate floras were limited (piasecki 1980). however, new comprehensive studies of dinoflagellate cysts throughout the danish neogene have resulted in a high-resolution strati graphy of the succession, even in marginal marine deposits where foraminifers and molluscs rarely occur. dinoflagellate cysts in strata of the central north sea basin were studied in offshore, hydrocarbon exploration wells, for comparison with the onshore, marginal, marine deposits and to obtain data from the youngest neogene that is not represented by onshore strata. the resulting stratigraphy is a so-called interval zonation, i.e. all zones are defined as successions of rocks between a lower and an upper event of fossil appearance or disappearance. for example the unipontidinium aquaeductum zone is defined from the first appearance of u. aquaeductum to the first appearance of achomosphaera andalousiense (figs 2, 3). however, the zone is also characterised by other associated events (stratigraphic first and last occurrences of other characteristic species) and by the dinoflagellate cyst assemblage in general. species with relatively common and consistent occurrence and with reported stratigraphic potential from outside the north sea basin, e.g. in the north atlantic realm, were preferentially selected for the definition of the zones. the zonation can therefore easily be compared with other international/north atlantic dinoflagellate zonations, despite the fact that much of the studied material was deposited in the relatively enclosed and possibly periodically brackish environments of the north sea basin. dinoflagellate cyst zonation the danish neogene is formally divided into 19 dinoflagellate cyst zones from the uppermost oligocene to the top of the pliocene, with one zone in the oligocene, 15 zones in the miocene and three zones in the pliocene (fig. 3). the pliocene is considered tripartite; zanclean, piacenzian and gelasian (see gradstein et al. 2004). the zonation has a high resolution in the uppermost oligocene, the lower miocene and the middle miocene (1, 8 and 4 zones respectively) and a lower resolution in the upper miocene and pliocene (3 and 3 zones respectively). the zones will be formally defined in a forthcoming paper. the proposed dinoflagellate cyst strati graphy for denmark correlates well with other formal stratigraphies in the north sea basin, e.g. in germany (köthe 2003), the netherlands (munsterman & brinkhuis 2004), belgium (e.g. louwye et al. 1999) and the united kingdom (e.g. head 1998), as well as with dinoflagellate cyst stratigraphies outside this basin in the north atlantic realm (de verteuil & norris 1996), in spite of differences in selection of diagnostic species. the new danish biozonation is also correlated with neogene chronostratigraphy and with nannoplankton biostratigraphy (see gradstein et al. 2004). geological results the new dinoflagellate cyst stratigraphy has been extensively tested in boreholes and exposures throughout jylland (fig. 1). the application provides a detailed stratigraphic framework for improved geological interpretations. three major, prograding deltaic sand systems with excellent potential as reservoirs for drinking water are recognised in the lower miocene to lowermost middle miocene in the central parts of jylland: the billund sand, the bastrup sand and the od derup formation (e.g. dybkjær 2004; rasmussen 2004). in addition, several problematic issues of neogene geology in the danish region have been solved: 1. the stratigraphic position of the vejle fjord formation (larsen & dinesen 1959) is now definitively established as latest oligocene (brejning clay) to earliest miocene (vejle fjord clay and sand) (dybkjær 2004; dybkjær & rasmussen 2007). 2. the upper, silty and sandy part of the sofienlund for mation is time equivalent to the vejle fjord clay and sand. the lower ulstrup clay and the sofienlund clay are time equivalent with the brejning clay (unpublished data 2006, k. dybkjær). 3. the oligocene–miocene boundary can be recognised in the eastern north sea basin as the last occurrence of common deflandrea phosphoritica (figs 2, 3). this event corresponds to the transition from the brejning clay to the vejle fjord 30 25 µm 25 µm a b c d 25 µm 25 µm fig. 2. stratigraphically significant dinoflagellate cysts from the miocene of denmark. a: unipontidinium aquaeductum. b: achomosphaera anda lousiense. c: chiropteridium galea. d: deflandrea phosphoritica. clay and to the sequence boundary b (rasmussen 2004; dybkjær & rasmussen 2007). 4. a significant hiatus is recognised between the vejle fjord and arnum formations in the northern and central parts of jylland. in the southern parts of jylland a time-equivalent fluvial-deltaic sand system, the ribe formation, was de posited (dybkjær & rasmussen 2000; dybkjær 2004; rasmussen 2004; rasmussen & dybkjær 2005). 5. the sand-rich succession in the salten profile and in nearby gravel pits (addit, voervadsbro) was formerly referred to the late early miocene to early middle miocene odderup formation. dinoflagellate cyst analysis has shown that this 31 5 10 15 20 b a nn21 nn20 nn19 nn18 nn17 nn16 nn15/ nn13 nn12 n n 11 nn10 nn9 nn8 nn6 nn5 nn4 nn3 nn2 nn1 np25 nn7 l l m m e e holocene pleistocene pl io ce ne m io ce ne oligocene a ge ( m a) epoch stage n an no pl an kt on zo na tio n dinoflagellate cysts zonation dinoflagellate events onshore zonation offshore zonation gelasian piacenzian zanclean messinian tortonian serravallian langhian burdigalian aquitanian chattian 23.03 20.43 15.97 13.65 11.61 7.25 5.33 3.60 2.59 1.81 amiculosphaera umbracula amiculosphaera umbracula impagidinium multiplexum bitectatodinium tepikiense melitasphaeridium choanophorum barssidinium pliocenium barssidinium evangelinae erymnodinium delectabile reticulosphaera actinocoronatum barssidinium evangelinae selenopemphix armageddonensis hystrichosphaeropsis obscura palaeocystodinium spp. gramocysta verricula achomosphaera andalousiense unipontidinium aquaeductum unipontidinium aquaeductum distatodinium biffii deflandrea phosphoritica, common chiropteridium galea caligodinium amiculum thalassiphora pelagica thalassiphora rota cordosphaeridium cantharellus exochosphaeridium insigne homotryblium spp. abundant labyrinthodinium truncatum cousteaudinium aubryae ectosphaeropsis burdigalensis exochosphaeridium insigne sumatradinium hamulatum palaeocystodinium miocaenicum h. obscura h. obscura s. armageddonensis m. choanophorum b. pliocenicum i. multiplexum g. verricula g. verricula a. andalousiense u. aquaeductum l. truncatum a. andalousiense a. umbraculaa. umbracula u. aquaeductum l. truncatum c. cantharellus c. cantharellus e. insigne e. insigne c. aubryae c. aubryae c. galea c. galea d. phosphoritica d. phosphoritica s. hamulatum s. hamulatum t. pelagica t. pelagica c. amiculum homotryblium spp. homotryblium spp. fig. 3. stratigraphic scheme presenting the new dinoflagellate cyst zonation correlated with the neogene nannoplankton zonation and chronostrati graphy. the time scale is according to gradstein et al. (2004). the diagnostic dinoflagellate cyst species are shown in red. 32 succession is older than previously assumed and is part of the vejle fjord formation/billund sand system (rasmussen et al. 2006). 6. the hodde transgression occurred in the early langhian, earliest middle miocene and limits the age of the underlying odderup formation to latest early miocene – earliest mid dle miocene (piasecki 2005). 7. the youngest onshore neogene deposits are of tortonian age (late miocene) as no messinian or pliocene dinoflagellate cysts have been recorded. the earlier sæd formation, upper gram formation sand and the neogene sand in cliffs at ho bugt are considered of tortonian age and referred to the gram formation (piasecki 2005). 8. a series of maps of the cenozoic succession in the danish north sea area was produced by rasmussen et al. (2005). the age determinations of the neogene succession are based on the dinoflagellate cyst stratigraphy presented here. future perspectives the new dinoflagellate cyst zonation is not only a valuable tool for unravelling the danish neogene succession. the neogene succession in denmark is presently being correlated with other neogene successions, e.g. in germany, the neth er lands and poland, in order to elucidate the neogene geology of the north sea basin. the zonation also leads to an improved understanding of the subsidence and tilting of the north sea basin during the neogene, and a better understanding of the petroleum systems in the hydrocarbon-producing provinces in the north sea area. this may lead to new discoveries and increased production from known oil/gas-fields (rasmussen et al. 2005). it is also an important tool for solving problems such as the evidence for and the timing of the neogene uplift of norway. acknowledgements the danish environment centres and the former counties are thanked for their cooperation and financial support. references de verteuil, l. & norris, g. 1996: miocene dinoflagellate stratigraphy and systematics of maryland and virginia. micropaleontology 42 (supplement), 172 pp. dybkjær, k. 2004: dinocyst stratigraphy and palynofacies studies used for refining a sequence stratigraphic model uppermost oligocene to lower miocene, jylland, denmark. review of palaeobotany and palynology 131, 201–249. dybkjær, k. & rasmussen, e.s. 2000: palynological dating of the oli gocene–miocene successions in the lille bælt area, denmark. bulletin of the geological society of denmark 47, 87–103. dybkjær, k. & rasmussen, e.s. 2007: dinocyst stratigraphy in an expanded oligocene–miocene boundary section in the eastern north sea basin (the frida-1 well, denmark) and correlation from basinal to marginal areas. journal of micropalaeontology 26, 1–17. fensome, r.a., riding, j.b. & taylor, f.j.r. 1996: chapter 6. dino flagellates. in: jansonius, j. & mcgregor, d.c. (eds): palynology: principles and applications. american association of stratigraphic pa ly nologists foundation 1, 107–169. gradstein, f.m., ogg, j.g. & smith, a.g. et al. 2004: a geological time scale 2004, 589 pp. cambridge: cambridge university press. head, m.j. 1998: marine environmental change in the pliocene and early pleistocene of eastern england; the dinoflagellate evidence reviewed. in: van kolfschoten, t. & gibbard, p.l. (eds) the dawn of the qua ternary. mededelingen nederlands instituut voor toegepaste geo wettenschappen tno 60, 199–226. köthe, a. 2003: dinozysten-zonierung im tertiär norddeutschland. re vue paléobiologie, geneve 22(2), 895–293. larsen, g. & dinesen, a. 1959: vejle fjord formationen ved brejning. sedimenterne og foraminiferfaunaen (oligocæn-miocæn). danmarks geologiske undersøgelse ii. række 82, 114 pp. louwye, s., de coninck, j. & verniers, j. 1999: dinoflagellate cyst strati graphy and depositional history of miocene and lower pliocene formations in northern belgium (southern north sea basin). geologie en mijnbouw 78, 31–46. munsterman, d.k. & brinkhuis, h. 2004: a southern north sea miocene dinoflagellate cyst zonation. netherlands journal of geosciences / geologie en mijnbouw 83, 267–285. piasecki, s. 1980: dinoflagellate cyst stratigraphy of the miocene hodde and gram formations, denmark. bulletin of the geological society of denmark 29, 53–76. piasecki, s. 2005: dinoflagellate cysts of the middle – upper miocene gram formation, denmark. in: roth, f. & hoedemarkers, k. (eds): the gram book. palaeontos 7, 29–45. rasmussen, e.s. 2004: stratigraphy and depositional evolution of the uppermost oligocene – miocene succession in western denmark. bulletin of the geological society of denmark 51, 89–109. rasmussen, e.s. & dybkjær, k. 2005: sequence stratigraphy of the upper oligocene – lower miocene of eastern jylland, denmark: role of structural relief and variable sediment supply in controlling sequence development. sedimentology 52, 25–63. rasmussen, e.s., dybkjær, k. & piasecki, s. 2004: the billund delta: a possible new giant aquifer in central and western jutland. geological survey of denmark and greenland bulletin 4, 21–24. rasmussen, e.s., piasecki, s., andsbjerg, j., dybkjær, k., vejbæk, o.v., jacobsen, c., britze, p. & bryde-auken, m., 2005: cenozoic maps of the danish north sea area. danmarks og grønlands geologiske under søgelse rapport 2005/33, 8 pp. rasmussen, e.s., dybkjær, k. & piasecki, s. 2006: neogene fluvial and nearshore marine deposits of the salten section, central jylland, denmark. bulletin of the geological society of denmark 53, 23–37. rasmussen, l.b. 1961: de miocæne formationer i danmark. danmarks geologiske undersøgelse iv række 4(5), 45 pp. zachos, j.c., pagani, m. sloan, l.c. thomas, e. & billups, k. 2001: trends, rhythms, and aberrations in global climate 65 ma to present. science 292, 686–693. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: kd@geus.dk geological survey of denmark and greenland bulletin 15, 2008, 45-48 the principal aim of mapping ground-water vulnerability in denmark is to ensure optimal protection of present and future drinking-water resources. groundwater vulnerability mapping of areas up to 1000 km2 has been taking place over the past seven years. the scale of mapping has been adjusted to meet the demands for details of regulation of land use requested by danish legislation. groundwater vulnerability mapping comprises analyses and integration of geological, geophysical, hydrological and geochemical data. this paper focuses on the geochemical reactions between groundwater and sediment. geochemical knowledge may sometimes not be fully and systematically utilised in groundwater vulnerability mapping. this paper presents different geochemical approaches and demonstrates how these can be successfully integrated with geological, geophysical and hydrological data. groundwater vulnerability mapping the national groundwater mapping project involves approximately 40% (17,476 km2) of the total danish land area classified as particularly valuable for groundwater abstraction, termed osd in danish. the mapping has been financed by a surcharge per cubic metre on consumed water. by 2007, approximately 17% (7066 km2) of denmark had been map ped (s. midby, personal communication 2008). groundwater mapping in denmark has focused on nitrate vulnerability mapping. the areas classified in 2005 as nitrate vulnerable osds, as based on pre-existing knowledge, or available intensive groundwater mapping, are shown in fig. 1. in 2005, approximately 15% of denmark had been delimitated as nitrate vulnerable osds according to the definitions given in danish legislation. the evaluation of nitrate vulnerability follows a series of criteria, and includes geochemistry of groundwater and aquifer, and the protecting capabilities of the overlying aquitards (miljøstyrelsen 2000). geochemical data types the method of large-scale groundwater vulnerability mapping differs from that of a process-oriented scientific field investigation, and it relies extensively on pre-existing data due to the high costs of acquiring new data with sufficient spatial density. the pre-existing data are located in the databases at the geo logical survey of denmark and greenland (geus) and have been collected over the last century. the geochemical data types used are point information from boreholes and consist of (1) groundwater chemistry data, (2) geochemical sediment data, and (3) colour descriptions of soil layers. the pre-existing geochemical data have been collected for other purposes than groundwater vulnerability mapping, e.g. drinking-water and irrigation wells. today, some of these wells have been closed because of poor water quality, but the data are still useful. the distribution in space and time of data from abstraction wells is uneven and these data therefore cannot stand alone. other existing data for vulnerability mapping come from the national groundwater monitoring network (stockmarr 2005). acquisition of new geochemical data in an osd area requires a detailed and carefully prepared strategy that presupposes a hydrogeological understanding of the area and an use of geochemistry in groundwater vulnerability mapping in denmark birgitte hansen and lærke thorling © geus, 2008. geological survey of denmark and greenland bulletin 15, 45–48. available at: www.geus.dk/publications/bull 45 50 km 53°n 55°n 57°n 8°e 10°e 12°e fig. 2a fig. 1. groundwater classification map showing abstraction areas that are vulnerable with respect to nitrate pollution of groundwater in denmark. modified from the region legislation plans from the county councils (amts råd) from 2005. evaluation of pre-existing geochemical data. new geochemical data are acquired by different techniques. the boreholes may be constructed by using the air-lift technique or conventional auger drilling. the wells are 5–200 m deep with 1–4 screens placed in the aquifer in order to cover the major groundwater quality variations. the advantage of both air-lift and auger drilling lies in the possibility of obtaining sediment samples for geological interpretation and geochemical analysis of the nitrate reduction capacity. the disadvantage is low resolution of the variations of the groundwater chemistry. the ellog auger drilling method (sørensen & larsen 1999), which allows contemporary groundwater sampling while drilling, provides very important data on the vertical distribution of the groundwater quality (thomsen et al. 2004). these data provide snapshots of the groundwater chemistry, because the techniques do not allow for permanent screens and soil sample collection. geochemical approaches groundwater chemistry the interpretation of groundwater quality must take the geochemical processes between water and sediment in the aquifer into account. each water analysis of redox-sensitive species reflects a certain redox state and can be used to evaluate nitrate vulnerability (miljøstyrelsen 2000). special attention must be paid to water samples showing signs of chemical disequilibrium between groundwater and sediment. such disequilibrium may indicate that nitrate is penetrating into the anaerobic part of the aquifer (thorling & thomsen 2001). the interpretation of the chemical evolution in time series is also an important tool for discovering significant trends towards more subaerobic conditions. the spatial density of groundwater quality data is often poor compared with that of other data types, as the sampling points (screens) are often placed in clusters and only some levels of the aquifers are represented. geochemistry of sediments the ability of the aquifer sediment to remove leached nitrate mainly depends on its concentration of reducing solid substances such as ferro-ions, pyrite and organic matter. therefore, another geochemical approach in vulnerability mapping is to estimate the nitrate-reduction capacity based on analyses of the sedimentary content of these substances (e.g. ernstsen et al. 2001). sediment samples should be obtained according to a detailed sampling plan that ensures that they are statistically representative of the examined geological succession, for example in agreement with the theory of sampling (e.g. petersen et al. 2005). it is important to obtain samples from both the aerobic and the anaerobic environment. if no detailed sampling plan has been worked out, we recommend collecting samples at one metre intervals in the borehole. the spatial density of the data points for determination of nitrate reduction capacity is often low compared to that obtained by means of other groundwater vulnerability mapping techniques, and this poses a general up-scaling problem (e.g. hansen et al. 2006). interpretation of the redox interface from soil colours the transition between aerobic sediment containing nitrate and anaerobic nitrate-free sediment is called the redox interface. geologically, this interface slowly moves downward in sandy layers by 0.01 mm to 10 cm per year as the reductants are oxidised by oxygen and nitrate (postma et al. 1991). the velocity of the downward movement of the redox interface depends on (1) the reactive content of nitrate-reducing matter; (2) the leaching of nitrate from land use; (3) the water flow and groundwater recharge; and (4) reaction kinetics. evaluation of the redox interface based on sedimentary colour descriptions is another feasible approach for nitrate vulnerability mapping. yellow, red and brown colours indi46 5 km 10°10′e10°e 56°10′n 56°05′n fig. 4 88.1346 redox interface (m b.s.) clay thickness <15 m osd areas30–100 15–30 7–15 0–7 0 redox interface (m b.s.) 75 50 25 0 80604020 pr ob ab ili ty ( % ) a b fig. 2. a: the thickness of clay layers and the position of the redox interface in an area south-west of århus, denmark. b: probability plot of the redoxinterface data shown in a. the location of the area is shown in fig. 1. cate aerobic conditions, while grey colours reflect anaerobic conditions. colour determination is subjective and is usually done in the field, often by using munsell soil colour charts. in areas with simple hydrogeological conditions, only one redox interface is found. complex hydrogeological conditions often result in several redox interfaces in the same borehole, which indicates non-vertical infiltration in a hete rogenous geological setting. the spatial density of the redox interface data is high compared with the density of groundwater chemistry data, and approaches the level of lithological data in an area. thus, the redox-interface data are crucial in groundwater nitrate vulnerability mapping. integration of geochemical and geophysical data in many areas, the thickness and composition of clay layers overlying aquifers play an important role in protecting groundwater against nitrate and other anthropogenic contaminants (thomsen et al. 2004). information about the nitrate vulnerability of an area can be obtained by combining data from geophysical mapping of the total clay content in the upper 30 m of the subsurface with area-distributed data of the redox interface based on soil colours. the geophysical data are commonly acquired using the pulled array continuous electrical sounding method (sørensen 1996). figure 2a shows an example from an approximately 150 km2 osd area in jylland, denmark. colour descriptions come from about 460 boreholes made in the period 1930–2006. the large variations in the depth of the redox interfaces, even between closely situated data points, are a consequence of the geological heterogeneity in the area. there is generally good agreement between areas where the redox interface has migrated to great depths (>7 m below surface: b.s.) and areas where the thickness of the clay layers is small (<15 m), presumably due to higher influx of oxidants because of high permeability. the depth of the redox interface in the area varies between the surface and c. 99 m b.s. in 50% of the investigated boreholes, the redox interface has migrated more than 7 m b.s. (the mean value in fig. 2b). these areas mostly correlate with areas of low groundwater protection due to low clay content. integration of geochemical, geological and hydrological data a borehole example appropriate mapping of groundwater nitrate vulnerability depends on integration of geochemical, geological, geophysical and hydrological data. figure 3 shows data from a 150 m deep borehole with four screens. data on water chemistry (columns 1–4), groundwater flow simulations based on the mike she programme (mainly horizontal from 40 m b.s., column 5), sediment chemistry (column 6), redox interfaces (column 7), and lithology (column 8) are presented. the quaternary deposits are dominated by sandy sediments with a composite thickness of about 100 m; with meltwater clay in the upper part and till beds in the lower part. the lowest part of the core penetrated impermeable eocene clay of the lillebælt clay formation. the colours of the sampled succession indicate seven redox interfaces in the borehole. the deepest one is at about 90 m b. s., and the aerobic/anaerobic zones correlate with the nitrate concentration in the groundwater. sedimentary pyrite concentration is very low in the anaerobic sandy aquifer (0.02–0.16 wt%), which indicates a very low nitrate reduction capacity (up to 40 years per m) of the aquifer and a relatively fast-moving redox interface (c. 2.5 cm per year). the simulated, mainly horizontal groundwater flow in the primary aquifer in the area of the borehole is very high (up to 4 m per year). the presence of nitrate and the low nitrate reduction capacity support these hydrological findings. lithological analyses of the concentration of different unstable ca-carbonate minerals in the fine gravel fraction show a high degree of weathering in the sandy meltwater deposits, although the groundwater remains saturated with calcite (hansen et al. 2006). these data also correlate well with the water chemistry (nitrate concentration), geochemistry of the sediments (low pyrite concentration), and ground water flow simulations (high groundwater flow). 47 1 oxygen mg/l 3 nitrite mg/l 4 sulphate mg/l 5 simulated flow mm/year 6 pyrite weight % 7 redox interfaces 8 lithology 0 0 4 00 2000 4000 0 240 80 .4 0 808 20 40 60 80 100 120 140 d ep th ( m b .s .) clayey till anaerobic layers aerobic layers clayey meltwater deposits eocene clay sandy meltwater deposits 2 nitrate mg/l fig. 3. chemical, hydrological and geological data for the 150 m deep well dgu no. 88.1346 located south-west of århus, denmark. the location of the well is shown in fig. 4. a profile example figure 4 shows an example of geological and geochemical conceptual models from an approximately 4 km long profile through a buried valley in jylland, denmark (e.g. jørgensen & sandersen 2008 – this volume). the left borehole (dgu no. 88.1346) on the profile is also shown in fig. 3. the boreholes and geophysical electrical data (see fig. 2) show that clay is of limited extent, which makes the aquifer highly vulnerable to nitrate and pesticides. the depth of the redox interface varies significantly along the profile and several redox interfaces can be found in the same borehole indicating a heterogeneous flow regime. nitrate, pesticides (des-dip-atrazine) and the pesticide degradation product bam have penetrated deeply, to about 90 m b.s. into the aquifer as shown by the yellow areas in fig. 4b. vast parts of the aquifer are polluted by nitrate and pesticides. the groundwater chemistry support the understanding of the redox interface and the geological setting. future extraction of groundwater from the deeper parts of the aquifer could worsen the nitrate and pesticide pollution of the primary aquifer due to increased flow. future perspectives in order to improve mapping of groundwater nitrate vulnerability there is a general need to develop the sampling and interpretation techniques for sedimentary geochemical an alysis. up-scaling of geochemical data interpretation in groundwater mapping is also a scientific field with high potential. the development of new methods for integrating geological, geophysical, hydrological and geochemical data is also very important. future focus should be on the development of guidelines for geochemical groundwater mapping and interpretation in a three dimensional perspective. the development and experience from denmark described here could be very useful for the eu member states working with large-scale mapping under the water framework directive. references ernstsen, v., henriksen, h.j. & von platen, f. 2001: principper for beregning af nitratreduktion i jordlagene under rodzonen, 54 pp. arbejdsrapport no. 24. københavn: miljøstyrelsen. hansen, b., jordt, b.e., thomsen, r., sørensen, j., kronborg, c. & nielsen, o.b. 2006: gebyrkortlægning i århus syd – geologisk, kemisk og hydro logisk datasammenstilling. geologisk nyt 3/06, 18–22. jørgensen, f. & sandersen, p.b.e. 2008: mapping of buried tunnel valleys in denmark: new perspectives for interpretation of the quaternary succession in denmark. geological survey of denmark and greenland bulletin 15, 33–36. miljøstyrelsen 2000: zonering. detailkortlægning af arealer til beskyttelse af grundvandsressourcen. vejledning fra miljøstyrelsen 3, 153 pp. køben havn: miljøstyrelsen. petersen, l., minkkinen, p. & esbensen, k.h. 2005: representative sampling for reliable data analysis: theory of sampling. chemometrics and intelligent laboratory systems 77, 262–277. postma, d., boesen, c., kristiansen, h. & larsen, f. 1991: nitrate reduction in an unconfined sandy aquifer: water chemistry, reduction processes, and geochemical modeling. water resources research 27, 2027–2045. sørensen, k. 1996: pulled array continuous electrical sounding. first break 14, 85–90. sørensen, k.i. & larsen, f. 1999: ellog auger drilling: 3-in-one method for hydrogeological data collection. ground water monitoring & remediation 19, 97–101. stockmarr, j. 2005: groundwater quality monitoring in denmark. geological survey of denmark and greenland bulletin 7, 33–36. thomsen, r., søndergaard, v.h. & sørensen, k.i. 2004: hydrogeological mapping as a basis for establishing site-specific groundwater protection zones in denmark. hydrogeology journal 12, 550–562. thorling, l. & thomsen, r. 2001: tunø. status report 1989–1999, 27 pp. århus: aarhus county, environmental division. authors’ address geological survey of denmark and greenland, lyseng allé 1, dk-8270 højbjerg, denmark. e-mail:bgh@geus.dk 48 anaerobic layers aerobic layers eocene clay quaternary clayey deposits quaternary sandy deposits potentiometric head redox interface sandy deposits (saalian/middle weichselian) clayey deposits (saalian/middle weichselian) clayey deposits (elsterian) sandy deposits (elsterian) eocene clay sandy deposits (elsterian/saalian/middle weichselian) clayey deposits (elsterian/saalian/middle weichselian) 88.1346 88.1344 88.529 nitrate: 50 mg/l des-dip-atrazine finds nitrate: 35 mg/l bam finds nitrate: 25 mg/l low nitrate reducing capacity high degree of calcite weathering high, simulated groundwater flow b a sea level 500 m 40 m w e w e sea level 40 m 500 m fig. 4. a buried valley south of århus, denmark. a: geological conceptual profile. b: geochemical conceptual profile. the boreholes are identified with dgu numbers. the location of the profile is shown in fig. 2a. geological survey of denmark and greenland bulletin 15, 2008, 37-40 over the years, several maps of the base quaternary surface of the danish area have been published. however, the maps have either been local in character (e.g. håkansson & pedersen 1992; huuse et al. 2001) or have concentrated on special topics such as tunnel valleys (e.g. huuse & lykke-andersen 2000) or glaciotectonic features (e.g. klint & pedersen 1995; and ersen et al. 2005). the only published map of a more regional character is that of binzer & stockmarr (1994) that covers onshore denmark and eastern danish waters. here we present for the first time a regional map of the base quaternary surface for the entire danish sector of the north sea and skagerrak based on interpretations of reflection seismic data at the geological survey of denmark and greenland (geus) (fig. 1). the new map has been depth-converted and merged with the onshore map of binzer & stockmarr (1994) and thus the first map covering the entire danish land and sea areas has been compiled. the definition of the base quaternary is a current issue of debate. in this article, we follow gradstein et al. (2004) who place the base quaternary at base gelasian, which is dated to 2.59 ma. in parts of the studied area, glacial tectonic features in the form of thrust complexes can be seen on the seismic data. here the base quaternary surface has been placed at the base of the dislocated thrust units, corresponding to the basal décollement horizon. the base quaternary surface is of both academic and practical interest. the depth to the base quaternary surface and its morphology are of interest to the understanding of base quaternary in the danish parts of the north sea and skagerrak tove nielsen, anders mathiesen and malene bryde-auken © geus, 2008. geological survey of denmark and greenland bulletin 15, 37–40. available at: www.geus.dk/publications/bull 37 two-way travel time (msec) 1200 0 10°e 10°e 56°n n 200 km 80 km 8°e4°e 8°e 56°n 57°n fig. 1. the base quaternary surface in the danish sector of the north sea and skagerrak. the depth to the surface is shown in seismic two-way travel time in milliseconds below sea surface. the inset shows the line density of the seismic database. the quaternary development of the region, but are also important in relation to offshore constructions such as oil and gas platforms, pipelines and wind mills. database and mapping procedure the present seismic interpretation of the base quaternary surface in the north sea and skagerrak (fig. 1) forms part of a larger north-west europe mapping project (petroleum geological atlas of the southern permian basin area – spba), and the mapping is based on geus’ digital 2d reflection seismic and well databases. the seismic interpretation is gridded in a 1000 × 1000 m grid and subsequently depth-converted using a linear time-depth function, ensuring that the resultant depth in metres fits available well data. in order to merge the new map with the pre-existing onshore map (binzer & stockmarr 1994), the latter has likewise been gridded into 1000 × 1000 m grids. to complete the composite map, a new seismic interpretation of the southern part of bælthavet has been carried out, following the same interpretation procedure as for the north sea and skagerrak areas. the base quaternary surface using the 2.59 ma age rather than the so far official 1.8 ma age for the base of the quaternary, only causes differences in the western north sea where the mapped surface is deeply buried and lies conformably on pliocene deltaic sediments 38 salt diapirs composite erosional unconformity neogene nne ssw sorgenfrei–tornquist zone norwegian trenchsalt diapir nesw chalk chalk 25 km conformable boundary transition zone erosional unconformity enewsw 500 1000 b c a faulted strata 500 1000 500 1000 25 km 25 km fig. 2. composite seismic profiles with the base quaternary surface marked in yellow. the location of the profiles is shown in fig. 3. a: cross-section of the danish north sea sector. in the eastern part, the base quaternary is seen as an erosional unconformity that cuts into midand late cenozoic delta deposits. towards the west, the base quaternary becomes conformable with the underlying lower pliocene delta clinoforms. b: cross-section of skagerrak where the quaternary deposits are relatively thin. the quaternary is underlain by chalk and marks a major hiatus most likely formed by erosion during multiple glaciations. the high-lying chalk area towards the south is coincident with the sorgenfrei–tornquist zone. towards the north, the base quaternary surface deepens and dips towards the norwegian trench. c: cross-section of the danish north sea sector illustrating how in large parts of the north sea the base quaternary marks a change in depositional style from prograding pre-quaternary delta deposits to aggradation or chaotic quaternary till deposits. towards the north, the shape of the base quaternary surface implies that in this area the salt domes have been active during the quaternary. (figs 1, 2). to the east, the base of the quaternary shallows and becomes an erosional unconformity that cuts into pre-quaternary deposits (figs 1, 2). the transition zone from the conformable boundary to the erosional surface is marked in fig. 3. the hiatus between the quaternary and the underlying strata becomes greater towards the east (rasmussen et al. 2005), and over large areas the base quaternary surface is underlain by chalk (fig. 2). this setting is a consequence of neogene uplift and subsequent erosion. the amount of glacial erosion during the quaternary has been estimated to about one fifth of the total neogene and quaternary erosion (japsen et al. 2002). the base quaternary erosional surface east of the transition zone is most likely to be a composite unconformity created during multiple glaciations. west of the transition zone, the lower quaternary deposits form a delta complex, and glacial erosion is only seen within the upper quaternary deposits in the form of tunnel valleys and internal erosional surfaces (fig. 2a). this depositional pattern shows that during the early quaternary the glaciations did not extend into the central north sea, and that full glaciation with grounding ice sheets did not occur in this area until later in the quaternary. in the north-eastern danish north sea, the base quater n ary surface displays an irregular morphology with several isolated, circular highs (figs 1, 3). in the westernmost part of the danish north sea three similar highs are also seen. this pattern is the result of salt doming that was active during the quaternary (fig. 2c). tunnel valleys occur in association with some of the salt domes, but few of them affect the base quaternary surface. in the south-western danish north sea, the base quaternary surface shows two elongated areas with depressions parallel to the coast of jylland (figs 1, 3). here glaciotectonic deformation has taken place, resulting in the formation of major thrust complexes involving prequaternary deposits with deep-lying décollement horizons. the trend of the thrust complexes and the internal pattern of the thrust faults indicate that they were formed by westwardmoving ice. in the central danish north sea, an e–w-trending belt of narrow, elongated incisions is seen in the base quaternary surface (figs 1, 3). this is a network of deep tunnel valleys that cuts into the pre-quaternary deposits. the valleys have the same overall n–s trend, suggesting that they were formed below ice advancing over the area from the north. some of the valleys are spatially associated with old fault systems (fig. 2a), but usually there appear to be no pre-quaternary features that can explain their location. it is worth noting that only few tunnel valleys incise the base quaternary surface outside this approximately 80 km wide belt (fig. 1), and hardly none on top of the ringkøbing–fyn high (fig. 3, inset). in most of the skagerrak area, the quaternary deposits rest unconformably on cretaceous chalk (figs 2b, 3). in the sorgenfrei–tornquist zone in southern skagerrak (fig. 3, inset), the base quaternary surface is relatively shallow due to up-thrusted chalk. towards the north, the surface deepens and dips in the direction of the norwegian trench (figs 1, 2). figure 4 shows the base quaternary composite depth map created by merging existing and new interpretations (see above). this map has been produced in order to place the interpretation of the north sea and skagerrak areas into a broader context. differences in databases and mapping procedure have caused some discrepancies in the composite map. nevertheless, some interesting observations can be made. the belt of distinct n–s-trending tunnel valleys incising the base quaternary surface in the north sea continues onshore in central jylland, where it is cut by a n–s-trending belt of shallower, e–w-trending tunnel valleys (figs 3, 4). the 39 fig. 2c fig . 2 b north sea kattegat skagerrak bælthavet western limit of chalk underlying quaternary deposits transition zone from conformable to erosive base quaternary seismic profiles international borders salt diapir province belt of north–south tunnel valleys glaciotectonic thrust complexes belt of east–west tunnel valleys 50 km 8°e 12°e 8°e 12°e 54°n 56°n 57°n ringkøbing–fyn high sorgenfrei–tornquist zone 100 km shallow basement fig. 2a jylland sjælland fyn fig. 3. overview map showing areas and outline of features mentioned in the text. also shown are the positions of the seismic profiles in fig. 2. the inset shows the main structural elements and outline of shallow basement. 40 e–w-trending tunnel valleys (pink on fig. 3) extend to the main stationary line (msl; the last glacial maximum stillstand line, e.g. houmark-nielsen 2007), indicating a late weichselian impact on the base quaternary surface in this onshore region, while the deeper n–s-trending tunnel valleys (light purple on fig. 3) lie south and west of the msl, suggesting a pre-weichselian age. acknowledgements fugro-geoteam and danpec a/s are thanked for permission to use their reprocessed 2d seismic data. references andersen, l.t., hansen, d.l. & huuse, m. 2005: numerical modelling of thrust structures in unconsolidated sediments: implications for glaciotectonic deformation. journal of structural geology 27, 587–596. binzer, k. & stockmarr, j. 1994: geological map of denmark, 1:500 000. pre-quaternary surface topography of denmark. danmarks geologiske undersøgelse kortserie 44, 10 pp., 2 maps. gradstein, f.m., ogg, j.g. & smith a.g. (eds) 2004: a geological time scale, 610 pp. cambridge: cambridge university press. håkansson, e. & pedersen, s.a.s. (eds) 1992: geologisk kort over den danske undergrund. varv. map sheet 1. københavn: varv. houmark-nielsen, m. 2007: extent and age of middle and late plei stocene glaciations and periglacial episodes in southern jylland, den mark. bulletin of the geological society of denmark 55, 9–35. huuse, m. & lykke-andersen, h. 2000: over-deepened quaternary valleys in the eastern danish north sea: morphology and origin. quat ernary science reviews 19, 1233–1253. huuse, m., lykke-andersen h. & michelsen, o. 2001: cenozoic evolution of the eastern danish north sea. marine geology 177, 243–269. japsen, p., bidstrup, t. & lidmar-bergström, k. 2002: neogene uplift of southern scandinavia induced by the rise of the south swedish dome. in: doré, a.g. et al. (eds): exhumation of circum-atlantic margins: timing mechanisms and implications for hydrocarbon exploration. geo logical society special publication (london) 196, 183–207. klint, k.e.s. & pedersen, s.a.s. 1995: the hanklit glaciotectonic thrust fault complex, mors, denmark. danmarks geologiske undersøgelse serie a 35, 30 pp. rasmussen, e.s., vejbæk, o.v., bidstrup, t., piasecki, s. & dybkjær, k. 2005: late cenozoic depositional history of the danish north sea basin: implications for the petroleum systems in the kraka, halfdan, siri and nini fields. in: doré, a.g. & vining, b.a. (eds): petroleum geology: north-west europe and global perspectives. proceedings of the 6th petroleum geology conference, 1347–1358. london: geological society. salomonsen, i. & jensen k.a. 1994: quaternary erosional surfaces in the danish north sea. boreas 23, 244–253. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: tn@geus.dk 5 5 5 5 5 56°n56°n 80 km 15°e 10°e5°e 15°e 5°e 57°n 55°n 56°n elevation of base quaternary surface 1230 m b.s.l. 150 m a.s.l. new interpretation existing map 200 km n fig. 4. base quaternary composite map. the map was created by merging the depth-converted new seismic mapping of the north sea, skagerrak and southern bælthavet with an existing map of the danish onshore, kattegat and inner domestic marine areas (binzer & stockmarr 1994) (see text for details). geological survey of denmark and greenland bulletin 15, 2008, 93-96 computer-controlled scanning electron microscopy (ccsem) combines the advantages of energy dispersive x-ray spectrometry (edx) with those of digital image analysis of back-scattered electron (bse) micrographs. ccsem analysis of a wide range of geological or non-geological materials has been introduced at the geological survey of denmark and greenland (geus) as a fast and reliable method to determine both the chemistry of individual grains and bulk samples. the chemical analysis is combined with measurements of the two-dimensional size and morphology of every single grain. the ccsem technique was developed in the early 1980s for characterisation of coal minerals (huggins et al. 1980; lee & kelly 1980) and studies of synthetic crystals for super-conductors and catalysts (lin & barnes 1984). soon it found a broader application in the study of dust particles and fibres in lung tissue of mine workers (friedrichs 1987), in the analyses of aerosols for air quality control and source emission characterisation (e.g. heasman & watt 1989) and the degree of sintering and consolidation of coal ash deposits (e.g. huffman et al. 1994). ccsem has been used in the earth sciences for the determination of the sediment budget of a lake (yin & johnson 1984), for the characterisation of soil and dust (pirrie et al. 2004), for provenance analysis of ilmenite-bearing beach sands (knudsen et al. 2005; bernstein et al. 2008), and provenance studies on sandstones in oil-bearing basins (frei et al. 2005). other areas where ccsem has been applied range widely and include characterisation of small inclusions, e.g. impurities in metal alloys or steel (schwoeble et al. 1988), analyses of gun-shot residues (e.g. steffen et al. 2007), and analyses of bladder stones obtained from a skeleton found in a mesolithic cave-tomb (d’alessio et al. 2005). in this paper, we demonstrate the benefits of the method with examples from the cement industry and from diamond prospecting. 93 fully automated analysis of grain chemistry, size and morphology by ccsem: examples from cement produc tion and diamond exploration nynke keulen, dirk frei, stefan bernstein, mark t. hutchison, christian knudsen and lucas jensen © geus, 2008. geological survey of denmark and greenland bulletin 15, 93–96. available at: www.geus.dk/publications/bull b a b 1 mm fig. 1. a: ccsem sample of beach sediment from jylland, denmark, divided into a number of frames in a grid. part of the grid is outlined in white. grains of different chemical compositions (different grey values) are embedded in epoxy resin. b: enlargement of one of the frames of the grid (indicated in red in a). the guard region (yellow) prevents double or incomplete measurements of grains (see text). the grey-level threshold function selects the grains one by one from the matrix for analyses of chemical composition, size and shape. analysed grains are shown in red; the image represents a snapshot of the ccsem procedure. analytical technique sample preparation sample material may, for example, consist of (1) a representative part of a bulk sample, (2) carefully selected grains mounted on double-sided tape, or (3) a heavy mineral separation of a bulk sample. grains, beads, and powders of both geological and non-geological origin can be analysed. for most studies, approximately 1 g of sample material was mounted in epoxy resin, using a technique that ensures that almost every grain is completely embedded in the epoxy, without touching any neighbouring grains (frei et al. 2005). the epoxy mounts are cut to show a representative part of the mount, subsequently polished and coated with carbon to enhance their conductivity. however, it is also possible to use thin sections of sample material prepared in a similar way. ccsem analysis the ccsem analysis was undertaken using a philips xl40 sem equipped with two edx detectors: a thermo nanotrace 30 mm2 window and a pioneer voyager 2.7 10 mm2 window si(li) detector. the tungsten filament of the sem was operated with an acceleration voltage of 17 kv, a filament current of typically 50–70 µa, and the sample was placed at a distance of 10 mm from the detector. the noran system six software package was used to automatically collect x-ray spectra, grain size and morphology of all particles and to recalculate the data following the proza (ϕρz) data correction and the filtering quantification technique. the technique described here is an improvement of the method described by frei et al. (2005) and bernstein et al. (2008). the samples were studied in the bse contrast mode of the electron microscope where the individual particles appear as different shades of grey in their black epoxy matrix (fig. 1). grey-level intensity thresholding by the image analysis function integrated in the software created a binary image of the bse micrograph and allowed for the separation and selection of individual grains (fig. 1b). a grid of image frames covering the whole sample area was defined by feeding the end-coordinates of the sample to the computer and by setting the required magnification (typically 30–100×) for the analysis (fig. 1a). grids consisted of 15 to 60 frames with approximately 20–35 grains per frame. a guard region between each frame ensured that double measurements of large particles in the sample was avoided and that only grains that lay completely within the image frame were included for analysis and thus recorded the true shape of grains. a ‘hole-fill’ function enabled more precise measurement of the grain size and shape from the binary image. because the grains were mounted in epoxy resin in such a way that they do not touch each other, no grain separation techniques had to be used, as commonly applied in automatic particle analysis software. thus, the original 2-d grain shape and grain size were completely available for analysis, without the introduction of artefacts by grain erosion and dilation or median filtering. all standard grain-shape factors can be measured. the smallest grains in the sample can be excluded from the analysis to avoid the measurement of particles that are only a few pixels in size, especially if a good grain morphology resolution is required. the created binary image formed the basis for the measurements of the grain chemistry. the software forced the microscope to scan within the perimeter of each grain to obtain the chemistry of either the whole grain area or from a single point in the centre of the grain mass. a typical spectrum for one particle contained 1000–2000 counts for the highest peak. spectra with a very low number of counts can be removed to ensure good measurement statistics. commonly, 800–1200 grains were measured in approximately two hours. the noran software produces a results table that lists shape, size and chemistry for each individual grain. all spectrum files and image frames, with a typical size of 1024 × 774 pixels, are stored after analysis. spectrum files can be reprocessed to include accidentally omitted elements retrospectively, without the need to physically reanalyse the sample. the chemical data are further reduced using a software package that is connected to a mineral library database for automatic phase recognition and data storage. 94 fig. 2. grain-size distribution diagram for seven minerals from raw materials used in the cement industry. grain mineralogy and grain size were determined with ccsem. 0 20 40 60 80 100 10 100 1000 feldspar bauxite hematite illites dolomite micas quartz c um ul at iv e w ei gh t (% ) grain size (µm) application of ccsem as a practical solution quality control of raw materials in cement production for cement production, raw materials are crushed to fine grain sizes before they are mixed and reacted at high temperatures in a kiln. to optimise the performance of the grinding mill and the sintering process an investigation of the grain-size distri bution of the limestone, iron ore, clay and sand fractions of the cement was undertaken. the grain-size distribution of the individual components in particular is of great importance for further processing of the raw material into cement clinker. the raw material was sieved into different grain-size fractions and treated with hydrochloric acid to decrease the amount of calcium carbonate in the raw material. the size fraction of 10–2000 µm was analysed with ccsem to determine the grain-size distribution and grain chemistry. figure 2 shows the grain-size distribution for seven different minerals. the mica, illite, and feldspar grains show uniform grain-size distributions, whereas bauxite, hematite and particularly dolomite grains show ranges of size distributions over two orders of magnitude. the hematite grains display a bimodal-size distribution. these findings demonstrate that ccsem is a suitable option for routine quality control and improvement of the production process. ccsem: a fast and reliable tool for diamond prospecting? determinations of the elemental composition of macro-crystalline phases in kimberlitic rocks or in detrital sediment samples are an important tool in diamond exploration. the major and minor element compositions of certain minerals are diagnostic for igneous rocks of mantle origin, and in some cases also represent a defined probability that the crystallisation of these phases occurred under conditions where diamond is stable. the ratios between cr and ca concentrations of mg-rich garnets and the cr and ca concentrations of eclogitic garnets are examples where such probability fields have been defined (fig. 3; grütter et al. 2004). a standard method for analysing the composition of garnets and other macrocrystalline phases is to measure the concentration of approximately ten to fifteen oxides with an electron microprobe (emp). emp is a dependable, yet time-consuming and relatively expensive method. we therefore tested the potential of ccsem analysis as a faster and cheaper method to measure the composition of macrocrystals in kimberlitic rocks. we used indicator minerals from the ‘garnet lake’ kimberlite body in west greenland, where diamonds are common (hutchison 2005). a series of hand-picked pyrope (garnet) grains were mounted in epoxy resin. the sample was analysed using ccsem, with extended counting times to ensure good statistics: the relative error in the reproducibility of the measurements is c. 1–2% for major elements and c. 4–8% for minor elements. the accuracy of ccsem was tested by comparing the results with compositional data obtained from electron microprobe analyses for the same minerals (as reported by hut ch ison 2005). even with extended counting time, a sample consisting of 200 grains can be analysed in 1–3 hours, with less than half an hour of operator time. an excellent reproduction of the emp measurements was achieved by ccsem (fig. 3); the statistical correlation between the two methods for these elements is 70%. the few outliers reflect those garnet grains that show a compositional gradient from core to rim. the emp point ana l yses were carried out on the cores of the grains, whereas the ccsem analyses average the whole surface of the grains, therefore providing slightly different results that are closer to the bulk composition of the grains. the vast majority of the garnet grains analysed by ccsem plot in the same classi fi cation field as do the grains according to the emp measurements. ccsem is therefore a good option for reliable and more ra pid measurements of garnet minerals in diamond-bearing rocks. validity of the ccsem measurements comparison between emp and ccsem shows high accuracy of the ccsem for minor elements. figure 4 shows the precision of the ccsem method for a major element (93.71 wt%), a minor element (2.19 wt%) and a trace element (0.21 wt%), measured on the same sample. five sets of measurements at nine different maximum peak count settings (equivalent to nine different time periods) were undertaken to evaluate the reproducibility of the data. for standard single spot or single grain analyses the relative errors are high com95 2 0 0 3 6 9 12 4 6 8 10 emp ccsem c r 2o 3 ( w t% ) cao (wt%) fig. 3. comparison of electron microprobe (emp) and ccsem measurements for chromium oxide and calcium oxide in mg-rich garnets from diamondbearing rocks. note the good correlation between the results obtained with ccsem and the emp data. pared to other analytical methods: 2–3% for major elements (>20 wt%), 5–10% for minor elements (>2 wt%) and 50–100% for elements present in smaller quantities. how ever, these figures can easily be improved by slightly increasing the counting time (fig. 4). the precision of bulk sample ccsem measurements is very high: as an effect of the long counting times the relative errors are reduced to <0.2% for major elements, <2% for minor elements and <15% for trace elements. this shows that the analysis time can be usefully tailored to the sample set depending on the required precision of the measurements and the time available. applied to diamond prospecting, measurement times could be adjusted depending on how marginal the sample is in the diamond stability field. the examples discussed above indicate that ccsem provides an accurate and precise way to rapidly measure single grain and bulk compositions of minerals and of other geological and non-geological materials. coupled with measurements of the grain size and other grain parameters for the individual particles, this technique is a potent tool to solve a wide range of problems. acknowledgement hudson resources inc. is thanked for provision of garnet samples. references bernstein, s., frei, d., mclimans, r.k., knudsen, c. & vasudev, v.n. 2008: application of ccsem to heavy mineral deposits: source of high-ti ilmenite sand deposits of south kerala beaches, sw india. journal of geochemical exploration 96, 25–42. d’alessio, a., bramanti, e., piperno, m. naccarato, g. vergamini, p. & forna ciari, g. 2005: an 8500-year-old bladder stone from uzzo cave (trapani): fourier transform-infrared spectroscopy. archaeomotry 47, 127–136. frei, d., rasmussen, t., knudsen, c., larsen, m., whitham, a. & morton, a. 2005: linking the faroese area and greenland: new methods and techniques used in an innovative, integrated provenance study. fróðskaparrit 43, 96–108. friedrichs, k.h. 1987: electron microscopic analyses of dust from the lungs and the lymph nodes of talc-mine employees. american industry hygiene asso ciation journal 48, 626–633. grütter, h.s., gurney, j.j., menzies, a.h. & winter, f. 2004: an updated classification scheme for mantle-derived garnet, for use by diamond explorers. lithos 77, 841–857. heasman i. & watt, j. 1989: particulate pollution case studies which illustrate uses of individual particle analysis by scanning electron microscopy. environmental geochemistry and health 11, 157–162. huffman, g.p. et al. 1994: investigation of ash by microscopic and spectroscopic techniques. in: williamson, j. & wigley, f. (eds): the impact of ash deposition on coal fired plants. proceedings of the engineering foun dation conference (20–25 june 1993, solihull, birmingham, uk), 409–423. london: taylor & francis. huggins, f.e., kosmack, d.a., huffman, g.p. & lee, r.j. 1980: coal mineralogy by sem analysis. scanning electron microscopy 1, 531–540. hutchison, m.t. 2005: diamondiferous kimberlites from the garnet lake area, west greenland: exploration, methodologies and petrochemistry. in: secher, k. & nielsen, m.n. (eds): workshop on greenland’s diamond po tential. 7–9 november 2005 in copenhagen. extended abstracts. dan marks og grønlands geologiske undersøgelse rapport 2005/68, 33–42. knudsen, c., frei, d., rasmussen, t., rasmussen, e. s. & mclimans, r. 2005: new methods in provenance studies based on heavy minerals: an example from miocene sands in jylland, denmark. geological survey of denmark and greenland bulletin 7, 29–32. lee, r.j. & kelly, j.f. 1980: overview of sem-based automated image analysis. scanning electron microscopy 1, 303 only. lin, m.c. & barnes, r.g. 1984: mössbauer spectroscopy and scanning electron microscopy study of iron-graphimet. journal of applied physics 55, 2294– 2296. pirrie, d., butcher, a.r., power, m.r., gottlieb, p. & miller, g.l. 2004: rapid quantitative mineral and phase analysis using automated scanning electron microscopy (qemscan); potential applications in forensic geoscience. in: pye, k. & croft, d.j. (eds): forensic geoscience: principles, techniques and applications. geological society special publication (london) 232, 123–136. schwoeble, a.j., dalley, a.m., henderson, b.c. & casuccio, g.s. 1988: com putercontrolled sem and microimaging of fine particles. journal of metals 40, 11–14. steffen, s., otto, m., niewoehner, l., barth, m., brozek-mucha, z., bieg straaten, j. & horvath, r. 2007: chemometric classification of gunshot residues based on energy dispersive x-ray microanalysis and inductively coupled plasma analysis with mass-spectrometric detection. spectrochimica acta b62, 1028– 1036. yin, c. & johnson, d.l. 1984: an individual particle analysis and budget study of onondaga lake sediments. limnology & oceanography 29, 1193–1201. 96 authors’ addresses n.k., d.f., m.t.h. & c.k., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ntk@geus.dk s.b., avannaa resources ltd., geological museum, university of copenhagen, øster voldgade 5–7, dk-1350 copenhagen k, denmark. l.j., flsmidth a/s, vigerslev allé 77, dk-2500 valby, denmark. 10 0 20 30 40 50 60 70 80 90 100 0 50 100 150 200 250 300 350 70006000 trace elements minor elements major elements time (s) r el at iv e er ro r (% ) bulk single grains fig. 4. precision of single grain and bulk sample ccsem analyses as a function of measurement times. two detectors were used for the analyses. the precision increases with counting time. geological survey of denmark and greenland bulletin 15, 2008, 89-92 reconstructing past secular environmental variations is an important issue in palaeoclimate research. however, most key variables for palaeoclimate reconstructions cannot be measured directly, and reconstructions are therefore based on proxy data. here, we demonstrate the potential of bivalve shells as an archive of environmental parameters. the geo lo gical survey of denmark and greenland (geus) has developed a fast and reliable method for chemical analyses of shell material by laser ablation inductively coupled plasma mass spectrometry (la-icp-ms), and here we present some ex amples of the use of this method. in tropical and subtropical waters, corals can provide century-long archives of past water chemistry with annual resolution. a comparable archive for temperate and arctic waters would be highly useful in climate research, and therefore it has been examined whether this can be provided by bivalve shells (e.g. schoene et al. 2005). long-lived species may provide archives with annual resolution extending over several hundred years, whereas short-lived, fast-growing species can provide archives with a seasonal or in some cases daily resolution over a period of a few years. most bivalves are sessile, and shells are commonly preserved as fossils. there are, however, a number of challenges related to the use of bivalves as proxy archives: (1) many proxies show species specific behaviour (seed 1980); (2) only very few proxies are dependent on a single variable (wefer et al. 1999); and (3) the effects of biology and ontogeny on the uptake of trace elements and stable isotope fractionation in shell carbonate are largely unknown and have to be evaluated empirically. therefore, any potential proxy must be calibrated individually for each species of interest before it can be used. a large number of chemical analyses are needed to calibrate a proxy. these are commonly obtained by solution icp-ms, in which sample preparation is time-consuming and labour-intensive. the use of la-icp-ms is therefore a considerable advance in bivalve shell proxy research, as it greatly reduces the effort needed for sample preparation. at the same time, the method requires less material for analysis, thus providing better spatial and hence temporal resolution. proxies based on bivalve shell carbonate can be used in present-day environmental monitoring, and for environmental reconstructions from shells found as fossils. shells from museum collections and shells found in archaeological middens can give information on historic and prehistoric environmental conditions (e.g. carrell et al. 1987), and fossil shells can be used as archives of environmental parameters on geological timescales (e.g. hendry et al. 2001). shell mineralisation bivalve shells consist mainly of calcium carbonate with im purities in the form of various elements substituting for calcium in the crystal structure. calcium carbonate represents 95–99% by weight of the shell, the remaining 1–5% being organic matrix, which is dominated by proteins (marin & luquet 2004). the shell material is deposited sequentially in growth increments that are often visible in polished sections studied by computer-controlled scanning electron micro scopy (ccsem). as a consequence of the growth pattern the increments occur in chronological order, and a relative time line for chemical analyses can be established. a section with the different layers of the shell of the common blue mussel (mytilus edulis) is shown in fig. 1, where internal growth increments are also illustrated. figure 2 shows an image of the actual shell structure. the elements needed for shell mineralisation come from the water or from particles that the bivalve ingests. in order to be included in the shell, the elements have to cross two biological membranes, the outer and inner mantle epithelium. these membranes actively discriminate against certain elements, but for some elements this discrimination is influenced by external stimuli (e.g. klein et al. 1996). 89 laser ablation analysis of bivalve shells – archives of environmental information maiken hansen klünder, dorothee hippler, rob witbaard and dirk frei © geus, 2008. geological survey of denmark and greenland bulletin 15, 89–92. available at: www.geus.dk/publications/bull aragonite calcite periostracum hemolymph m an tle sh el l inner extrapallial fluid (epf) outer extrapallial fluid epithelium (inner/outer) fig. 1. section through the margin of shell and mantle of a mytilus edulis. the crystalline shell consists of two separate layers: a prismatic layer of calcite and an aragonitic layer of nacre. the outermost layer is a protective organic layer (periostracum). the shell is secreted in growth increments in the area be tween the shell and the mantle. the la-icp-ms method many previous studies of bivalve shells have utilised wet chemical analysis. samples often consist of powder drilled from the shell with dentist drills and other microdrilling tools. powder samples are routinely analysed by solution icp-ms. la-icp-ms combines an analytical precision comparable to that of solution icp-ms with a significantly shorter and easier sample preparation process. the laser technique is not only time-saving – the fewer steps needed in sample preparation also reduce the risk of contamination. furthermore, the spatial resolution is much higher, as laser ablation in shell samples can be undertaken with a beam diameter of 30–65 µm, as opposed to the 200–300 µm diameter of a microdrill. sample preparation any sample of bivalve shell material can be analysed by laicp-ms, but cross-sections through entire valves are preferred in order to constrain the growth history. the shell must be cleaned of soft tissues, epibionts or adhering sediment. the shell material is embedded in epoxy resin to prevent it from fracturing during handling. the shell is then cut with a diamond-tipped rock saw to produce a cross-section, and polished to show the shell structure (fig. 3). shells longer than 5 cm may have to be divided into two or more sections to fit into the sample chamber of commercially available laser ablation systems. after polishing, the section is cleaned with alcohol and treated ultrasonically to remove possible surface contamination. analytical techniques the la-icp-ms equipment at geus is a finnigan el ement2 high resolution icp-ms connected to a new wave research up213 laser ablation system. for shell analyses, the nist 612 and nist 614 glasses are used as standard materials. the elemental concentrations for the standard glasses published by pearce et al. (1997) are used for concentration calculations. there are potential problems in using nonmatrix-matched standards, but at ablation times of less than 80 seconds, these problems are not significant in analyses carried out on calcium carbonate (vander putten et al. 1999). as an internal standard in the samples, calcium (43ca) is suitable for the measurement of several trace and minor elements in calcite (longerich et al. 1996), and sem analyses of m. edulis have shown that the calcium content in the calcite layer is uniform. we use the glitter software package for final concentration calculations from the time-resolved raw data. relative age and growth rate an advantage of calibrating a proxy on bivalve shells taken from laboratory or field culturing experiments is that measurements of the shell length can be made during the experiment, so that the chemical analyses can be time constrained. when applying the proxy to fossil shells, it is of course im possible to carry out multiple shell length measurements on the live individual, so other methods must be used. many species form annual growth increments that can be used to set the relative age of a specimen. furthermore, all species show micro-increments that are visible in a microscope. these narrow growth increments are not always regular, but in a number of species the increments show a periodicity related to moon phases or diurnal or tidal shifts. utilising what is known about the periodicity of increment formation in the species analysed, one can assign relative ages to chemical analyses and calculate approximate growth rates for the analysed shell. examples mg/ca thermometry the use of the ratio between mg and ca in shells as a temperature proxy was first suggested because it was found that the mg/ca ratio in marine carbonates varies according to lat90 5 µm 1 cm fig. 2. sem image of a mytilus edulis shell in cross-section, showing aragonitic nacre (left side of image) and prismatic calcite (right side of image). the image illustrates the differences in structure between these two shell layers, and the direction of the growth increments. fig. 3. computer scan of shell sample mytilus edulis b218 prepared for laicp-ms analysis. itude (see henderson 2002). using calcite from m. edulis taken from field culturing experiments in the wadden sea, we found shell mg/ca ratios to be temperature dependent. the mg/ca ratio of m. edulis shells from svendborg sund, denmark was then used to calculate seawater temperatures. the temperature was calculated from the mg/ca ratio using the equation t = 2.22 + 18.2 log (mg/ca) (unpublished data, m.h. klünder). the calculated temperatures have been compared with water temperatures measured by the national en vironmental research institute (fig. 4). it is seen that the mg/ca thermometer gives a fair estimate of temperature changes during a summer. lead pollution the concentration of pb in shell increments of the bivalve mya arenaria is a function of the pb concentration in the water (pitts & wallace 1994), and hence the former concentration of pb in the water can be calculated from the lead concentration in m. arenaria shells. shell samples from a fouryear old specimen from limfjorden, denmark, collected in 2005, have been analysed. the results indicate that the pb concentration in limfjorden has varied from 20 to 280 pmol/kg water over the sampled time span (fig. 5). hence analysing a single water sample may give a misleading picture of the pb level. the concentration of pb in limfjorden is comparable to that found in the relatively uncontaminated cape cod bay, eastern usa; it is up to ten times higher than pre-industrial levels in the boston area, as calculated from the pb content of sub-fossil shells from shell middens, and ten times lower than in boston harbour (pitts & wallace 1994). the pb proxy has also been applied to data from an arctica islandica individual that was transferred from the baltic sea to a dutch harbour (fig. 5). the proxy has not yet been calibrated for a. islandica, and the results can only be regarded as qualitative. however, it is seen that the pb concentration in shell material secreted after transplantation to the harbour is significantly higher than in that secreted in the baltic. these results indicate that shells of a. islandica can be used to monitor pb contamination of seawater. shell mn/ca and ba/ca – a link to primary production? it has been suggested that the content of mn and/or ba in bivalve shells can be correlated with primary production (e.g. stecher et al. 1996). this would suggest that mn/ca or ba/ca ratios are related to phytoplankton blooms, providing a proxy 91 0 5 10 15 20 25 0 50 100 150 200 jan/1 mar/1 may/1 jul/1 sep/1 nov/1 jan/1 calculated temperature measured temperature t em pe ra tu re ( °c ) analysis number 0 100 200 300 400 020406080100120140160 0 10 20 30 40 50 60 pm ol p b/ kg w at er analysis no. a545 analysis no. 025 10 0 20 30 40 50 c hl -a , n on -a ci di fie d μm ol b a/ m ol c a 0.5 1 1.5 2 2.5 3 3.5 4 10 0 20 30 40 50 0.05 0.10 0 0.15 sep/1 jan/1 may/1 sep/1 jan/1 c hl -a , n on -a ci di fie d μm ol m n/ m ol c a a b fig. 4. temperatures in svendborg sund, denmark in the summer of 2005 calculated from mg/ca ratios in mytilus edulis (red line) and compared to measured water temperatures (blue line). the calculated temperatures provide a fair estimate of the actual temperature. fig. 5. pb content in seawater calculated from shell pb concentration using the equation of pitts & wallace (1994). one specimen of mya arenaria (blue line, limfjorden) and one specimen of arctica islandica (red line, baltic sea, transplanted to the netherlands) were analysed. the arrow shows the time of the transplantation. fig. 6. ba/ca and mn/ca ratios of four mytilus edulis specimens compared with the chlorophyll-a concentrations (red line) of the seawater. a: ba/ca ratios in m. edulis shells from the dutch wadden sea. b: mn/ca measured in the same shells. for the timing and size of such events. to test this theory, m. edulis shell samples from an aquaculture field experiment site in the wadden sea were analysed at geus and compared to the chlorophyll-a concentration of the ambient water (fig. 6). the results are not conclusive, but they suggest that the relationships between the mn/ca or ba/ca ratios and the chlorophyll-a concentration are not simple linear functions. the ba/ca ratio in the shells seems to increase with the chlorophyll-a concentration in the water, but continues to remain at an elevated level after the end of the bloom; however, mn/ca seems to have a correspondence with the peaks of some less intensive algae blooms, but is quite low during the most pronounced bloom recorded in may. clearly, further research is needed to better understand the link between mn/ca and ba/ca ratios in shell material and phytoplankton blooms. final remarks the la-icp-ms method is a reliable and advantageous technique for the analysis of a wide range of trace elements in carbonates. the combination of relatively high precision, low detection limits, high spatial resolution, straightforward sample preparation and fast analysis makes the method especially suited for research and application of calcium carbonate based proxies. further development of biogenic carbonate proxies will have benefits for both palaeo-climate research and investigation into the processes of biomineralisation, as well as for environmental studies. acknowledgements we would like to thank m. sejr for assistance and j. v. iperen for the chlorophyll-a data. this paper is a contribution to the euroclimate project 04 eclim fp08 casiopeia. the work has been supported financially by the european science foundation under the eurocores programme euroclimate and the danish research agency for science, technology and innovation (fnu). references carell, b., forberg, s., grundelius, e., henrikson, l., johnels, a., lindh, u., mutvei, h., olsson, m., svärtström, k. & westermark, t. 1987: can mussel shells reveal environmental history? ambio 16, 2–10. henderson, g.m. 2002: new oceanic proxies for paleoclimate. earth and planetary science letters 203, 1–13. hendry, j.p., perkins, w.t. & bane, t. 2001: short-term environmental change in a jurassic lagoon deduced from geochemical trends in aragonite bivalve shells. geological society of america bulletin 113, 790–798. klein, r.t., lohmann, k.c. & thayer, c.w. 1996: bivalve skeletons record sea-surface temperature and δ18o via mg/ca and 18o/16o ratios. geology 24, 415–418. longerich, h.p., günther, d. & jackson, s.e. 1996: elemental fractionation in laser ablation inductively coupled plasma mass spectrometry. fresenius’ journal of analytical chemistry 355, 538–542. marin, f. & luquet, g. 2004: molluscan shell proteins. comptes rendus palevol 3, 469–492. pearce, n.j.g., perkins, w.t., westgate, j.a., gorton, m.p., jackson, s.e., neal, c.r. & chenery, s.p. 1997: a compilation of new and published major and trace element data for nist srm 610 and nist srm 612 glass reference materials. geostandards newsletter 21, 115–144. pitts, l.c. & wallace, g.t. 1994: lead deposition in the shell of the bivalve mya arenaria: an indicator of dissolved lead in seawater. estuarine, coastal and shelf science 39, 93–104. schoene, b.r., fiebig, j., pfeifer, m., gless, r., hickson, j., johnson, a.l.a., dreyer, w. & oschmann, w. 2005: climate record from a bivalved methuselah (arctica islandica, mollusca, iceland). palaeo geography, palaeoecology, palaeoclimatology 228, 130–148. seed, r. 1980: shell growth and form in the bivalvia. in: rhoads, d.c. & lutz, r.a. (eds): skeletal growth of aquatic organisms, biological records of environmental change, 23–67. new york: plenum press. stecher, h.a., krantz, d.e., lord, c.j., luther, g.w. & bock, k.w. 1996: profiles of strontium and barium in mercenaria mercenaria and spisula solidissima shells. geochimica et cosmochimica acta 60, 3445–3456. vander putten, e., dehairs, f., andré, l. & bayens, w. 1999: quantitative in situ microanalysis of minor and trace elements in biogenic calcite using infrared laser ablation – inductively coupled plasma mass spectrometry: a critical evaluation. analytica chimica acta 378, 261–272. wefer, g., berger, w.h., bijma, j. & fischer, g. 1999: clues to ocean history: a brief overview of proxies. in: fischer, g. & wefer, g. (eds): use of proxies in paleoceanography: examples from the south atlantic, 1–68. berlin: springer. 92 authors’ addresses m.h.k. & d.f., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: mhk@geus.dk d.h., faculty of earth and life sciences, vrije universiteit amsterdam, de boelelaan 1085, 1081 hv amsterdam, the netherlands. r.w., royal netherlands institute for sea research (nioz), p.o. box 59, 1790 ab den burg (texel), the netherlands. geological survey of denmark and greenland bulletin 15, 2008, 49-52 this paper describes an ongoing multidisciplinary study on the development of the barrier islands in the danish wadden sea (vadehavet), carried out by the department of geography and geology at the university of copenhagen and the geolo gical survey of denmark and greenland (geus). nine sediment cores each c. 25 m long and a total of c. 45 km ground penetrating radar (gpr) profiles have been acquired on the islands of rømø and fanø. geochemical and palaeontological analyses and dating of 150 core samples using optically stimulated luminescence (osl) are in progress. this multidisciplinary approach has given new insights into the sedimentary architecture and development of the island, and the study is expected to result in a new detailed facies model. such models are essential for an assessment of the effects of rising sea level associated with global warming. the new facies model can also be used as an analogue for subsurface oil or water reservoirs in similar sedimentary settings. this article presents selected core and gpr data from the rømø barrier island. setting the rømø barrier island is situated in the northern part of the european wadden sea (fig. 1). the maximum tidal amplitude is c 1.8 m (andersen & pejrup 2001). during the last c. 8000 years, the area has experienced an overall relative sea-level rise of ~15 m (behre 2007). the island of rømø is c. 14 km long and c. 4 km wide and separated from the mainland by a c. 8 km wide lagoon (figs 1, 2). the island is connected to the mainland by a dam. tidal inlets, c. 1 km across, occur at the northern and southern tips of the island. the inlets continue as tidal channels into the lagoon and cut sand flats and mixed flats. subtidal ebb-deltas are located where the tidal inlets terminate in the north sea (fig. 3). the inlets reach depths of up to 30 m, as in the listerdyb south of rømø, or 4–10 m in the juvre dyb north of the island. the salt marsh area fringing the lagoonal coast of the island is up to 2 km broad. sand flats, up to 2.7 km wide, with three, 0.8–1.6 km broad bars that migrate towards the island, characterise the north-western and southwestern parts of the island (fig. 2). tidal creeks, up to c. 50 m wide, separate the bars. active eastward migrating aeolian dunes are found on large parts of the island. sedimentary facies of the rømø barrier island core wells seven core wells were drilled on the island, and depositional units were defined on the basis of sedimentary structures, grain sizes, sorting, organic material, fossils, trace fossils and rootlets. data from the rømø-4 and -1 wells are presented here. © geus, 2008. geological survey of denmark and greenland bulletin 15, 49–52. available at: www.geus.dk/publications/bull sedimentary facies and architecture of the holocene to recent rømø barrier island in the danish wadden sea peter n. johannessen, lars henrik nielsen, lars nielsen, ingelise møller, morten pejrup, thorbjørn j. andersen, joakim korshøj, birger larsen and stefan piasecki 49 salt marsh mud flat land, supratidal river mixed mud flat sand flat tidal channels more than 6 m deep subtidal 5 km 55°n 9°e 9°e denmark germany north sea 250 km fanø rømø sylt mandø juvre dyb listerdyb north sea wadden sea fig. 1. the danish wadden sea. the distribution of sediment types and subtidal channels are shown. the mainland of jylland east of the lagoon consists of glacial deposits. inset map shows the whole tidal wadden sea area. modified from pejrup (2006) and sørensen et al. (2006). in the rømø-4 well the upper c. 3.5 m consist of wellsorted aeolian sand (fig. 4). a mud bed, c. 20 cm thick and supposed to be deposited in a depression (swale) underlies the aeolian sand. about 11 m of medium-grained sand with numerous bivalve shells underlie the mud layer and are interpreted to have been deposited in a prograding to aggrading marine shoreface. at c. 2 m below mean s.l., a c. 0.3 m thick layer of coarse-grained sand may represent washover fans. the rømø-1 well that was drilled in the lagoon east of the island shows c. 8 m of bioturbated heteroliths of sand and mud with numerous bivalve shells, especially mytilus, and gastropods below the dam (figs 2, 4). the heteroliths represent back-barrier lagoonal mudand sand flats and overlie 3 m of sand-streaked mud with numerous bivalve shells. gyttja and peat layers, respectively 20 and 25 cm thick, underlie the mud (fig. 4). the peat layer was probably deposited during the initial holocene sea-level rise. continued rise in sea level caused flooding of the peat swamp and formation of gytja, followed by lagoonal mudand sand flats. the holocene sediments overlie pleistocene and eeminan deposits. the variation in composition of the dinoflagellate assemblages and other fossil algae from the core samples reflects shallow marine to lagoonal environments, in part with lowered salinity. the fauna of bivalves, snails and sea urchins in the cores is similar to that seen on the beaches of rømø today. the rich but low-diversity fauna seen in the lagoonal deposit in the rømø-1 well differs from the more diverse fauna with thick shells that is washed ashore on the open western coast at present. the fauna in the rømø-4 well consists of small shells of the west coast fauna only. this suggests that the sand was deposited by washover events from the north sea and that larger shells were left behind. the interpretations of the palynomorph assemblages and the macrofauna support the sedimentological interpretations and will be used in the reconstruction of the physical development of rømø. ground penetrating radar (gpr) w–e and n–s-trending gpr sections with a total length of c. 30 km were acquired from rømø using unshielded 100 mhz antennae manufactured by sensors & software inc. 50 n 6 t07a-1 1 km 4 7 5 3 2 1 fig. 2. orthophoto of the rømø barrier island. the location of gpr reflection profiles and the seven core wells are shown. very wide tidal sand flat characterise the north-western and south-western parts of rømø. rømø island is dominated by aeolian dunes that are migrating eastwards. copyright scankort. skallingen fanø mandø north sea rømø 10 km n fig. 3. landsat image (+etm 9 may 2001) of the danish wadden sea with distinct tidal sand flats and subtidal channels. note the subtidal ebbdeltas where the main channels terminate in the north sea. (fig. 2; nielsen et al. in press). maximum signal penetration is c. 15 m in the central parts of the island where fresh groundwater is thickest, and the vertical resolution is 0.2–0.3 m. the reflected signal only reaches about 1 m depth at the margins of rømø because of strong damping of the electromagnetic waves due to saltwater intrusion. salt marsh, peat and mud layers also significantly reduce the signal penetration. a gpr section from the central part of the island shows a beach ridge, c. 1.25 m high and c. 70 m wide, with a steep erosional side towards the west and a less steep side towards the east (fig. 5). the top of the beach ridge is c. 0.5 m above mean s.l. two superimposed cross-bedded units with eastdipping foresets occur upon the beach ridge. the lower unit is 0.8 m thick and c. 30 m long and the upper unit is c. 1 m thick and c. 50 m long (fig. 5). east of the washover fans large, gently eastward-dipping clinoforms occur with much higher amplitude signal, which indicates that the sediments are organic or mud rich. data from the rømø-4 core well situated c. 75 m north of the section indicate that the high amplitude layer correlates with a 20 cm thick mud layer deposited in a swale (fig. 4), consistent with previous findings elsewhere on the island (nielsen et al. in press). a more than 125 m wide set with clinoforms, c. 1.5 m thick, dipping in an eastward direction is found in the westernmost part of the section (fig. 5). the set terminates eastwards in a channel-like structure. the clinoform set is interpreted as an eastward-migrating tidal bar filling in the western side of a 1.5 m deep tidal channel, forcing the channel eastwards, which is comparable to the modern situation illustrated in fig. 3. the amplitude of the reflections in the channel sediments differs from those of the bar, suggesting that the channel fill is slightly muddier due to channel abandonment. the top of the channel and the bar are situated at the modern mean s.l. 51 rømø-4 clay si vf f m sand c vc peb. clay si vf f m sand c vc peb. prograding/agrading shoreface sand prograding upper shoreface sand aeolian sand washover fans? swale h ol oc en e w ei ch se lia n w ei ch se lia n h ol oc en e –1 –2 –3 –4 –5 –6 –7 –8 –9 –10 –11 –12 marine transgressive surface metre 4 3 2 1 0 metre rømø-1 dam back-barrier lagoon with mud and sand flats fluvial channels lagoon lagoonal transgressive surface peat mytilus mytilus mytilus mytilus gyttja –11 –10 –12 –13 1.9 1 0.0 –1 –2 –3 –4 –5 –6 –7 –8 –9 legend lithology structures fossils biogenic structure peat clay heterolith 80% clay, 20% sand sand silty sand pebbles clay clast cross-bedding parallel lamination current ripple cross-lamination rootlets shells sea urchin organic rich clast bioturbation gastropods fig. 4. sedimentological core logs from the rømø-1 and -4 wells, situated in the lagoon and the centre of rømø, respectively. for locations see fig. 2. 52 between the bar-channel complex and the beach ridge occur two sediment wedges, altogether c. 150 m across (fig. 5). the eastern wedge onlaps the relatively steep western side of the beach ridge, whereas the western wedge onlaps the eastern wedge and is truncated by the channel to the west. the wedges were probably deposited by landward migrating bars that welded to the coast, causing a stepwise shoreface progradation. aeolian sand lies on top of these foreshore sediments (fig. 5). the gently westward dipping, high amplitude reflections in the deeper part of the section are interpreted as representing westward shoreface progradation, in agreement with the core data (figs 4, 5). discussion and conclusions the cores and gpr sections provide high quality data for the identification of the depositional units that compose the rømø barrier island, and a few examples of the units are presented here. the interpretations of the cores and the gpr sections show that the rømø barrier island was never located as far east as the rømø-1 well. furthermore, lagoonal sediments are not present at the position of the rømø-4 well suggesting that lagoonal conditions were never established at this location. however, data from the cores of the rømø-6 and -3 wells (to be described in forthcoming papers) suggest that the western fringe of the lagoon reached the positions of these wells, because lagoonal deposits are identified below aeolian sand (nielsen et al. in press). samples of sand from the cores have been submitted for osl dating, and shells and organic matter are being dated by the 14c method. when the ages of the depositional units become available the development of the depositional units through time will be described in detail. plots of sample depths against ages will be constructed to portray the relative sea level history and will, together with analyses of the mollusc fauna and palynomorph assemblages, provide additional constraints on the reconstructions of the development of the barrier island. the study will be concluded with a detailed reconstruction of the holocene development of the rømø barrier island. acknowledgements the danish natural science research council and geocenter copenhagen are thanked for financial support. references andersen, t.j. & pejrup, m. 2001. suspended sediment transport on a temperate, microtidal mudflat, the danish wadden sea. marine geology 173, 69–85. behre, k.-e. 2007: a new holocene sea-level curve for the southern north sea. boreas 36, 82–102. nielsen, l., møller, i., nielsen, l.h., johannessen, p.n., pejrup, m., korshøj, j.s. & andersen, t.j. in press: integrated ground-penetrating radar and sedimentological studies of wadden sea barrier islands. journal of applied geophysics. pejrup, m. 2006: det danske vadehav og store landvindingsprojekter. in: sand-jensen, k. & fenchel, t. (eds): naturen i danmark: havet, 433–415. københavn: gyldendal. sørensen, t.h., bartholdy, j. christiansen, c. & pedersen, j.b.t. 2006: intertidal surface type mapping in the danish wadden sea. marine geology 235, 87–99. authors’ addresses p.n.j., l.h.n., i.m., b.l. & s.p.: geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: pjo@geus.dk l.n., m.p., t.j.a. & j.k.: department of geography and geology, øster voldgade 10, dk-1350 copenhagen k, denmark. wofwof aeoaeo aeoaeo sfclsfcl swsw chch chch brbr wofwofwof b b b 5 0 aeo aeo sfcl sw ch ch br wof el ev at io n (m ) –5 w e 100 m b b b rømø-4 well c. 75 m north of gpr profile aeo: aeolian, br: beach ridge, sw: swale, wof: washover fan, b: bar, ch: channel, sfcl: shoreface clinoforms : present mean sea level fig. 5. west–east-trending ground penetrating radar profile t07a-1 from rømø. the upper part of the section consists of aeolian sand. for location see fig. 2. geological survey of denmark and greenland bulletin 27, 2012, pp. 68 1 geological survey of denmark and greenland bulletin 27 • 2012 neoglacial and historical glacier changes around kangersuneq fjord in southern west greenland anker weidick, ole bennike, michele citterio and niels nørgaard-pedersen geological survey of denmark and greenland danish ministry of climate, energy and building 22 geological survey of denmark and greenland bulletin 27 keywords greenland, godthåbsfjord, kangersuneq, kangiata nunaata sermia, glaciology, tidewater glaciers, quaternary, holocene, neoglacial, little ice age. cover the calving front of kangiata nunaata sermia seen from the south-west. the trimline zone on nunataarsuk semi-nunatak in the background reaches almost to the top of the mountains. the calving front is 4.5 km long. photograph: dirk van as, 22 august 2011. frontispiece: facing page qamanaarsuup sermia is separated into an upper and a lower part by a bedrock high. the area in the foreground was covered by an ice-dammed lake during the little ice age maximum. the lower part of the glacier is c. 2 km wide. photograph: dirk van as, 28 august 2012. chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, department of geoscience, aarhus university; minik rosing, geological museum, university of copenhagen; finn surlyk, department of geography and geology, university of copenhagen scientific editor of this volume: jon r. ineson editorial secretaries: jane holst and esben w. glendal referees: ole humlum (no) and niels tvis knudsen (dk) illustrations: jette halskov digital photographic work: benny m. schark layout and graphic production: kristian a. rasmussen printers: rosendahls schultz grafisk a/s, albertslund, denmark manuscript received: 31 january 2012 final version approved: 24 august 2012 printed: 27 december 2012 issn 1604-8156 isbn 978-87-7871-347-6 citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 27, 68 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk or at www.geus.dk/publications/bull © de nationale geologiske undersøgelser for danmark og grønland (geus), 2012 for the full text of the geus copyright clause, please refer to www.geus.dk/publications/bull mailto:geus@geus.dk www.geus.dk/publications/bull www.geus.dk/publications/bull 3 44 contents abstract . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5 introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6 setting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7 notes on the holocene history of the area . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9 geological, archaeological and historical information on glacier fluctuations and sea-level changes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 14 data from aerial photographs and satellite images . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 15 glacier changes in the kangersuneq area . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17 the former ice-dammed lake isvand . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17 evaluation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 22 the kangiata nunaata sermia glacier system . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23 evaluation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 29 qamanaarsuup sermia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 31 evaluation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 32 narsap sermia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 32 evaluation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 33 kangilinnguata sermia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 35 evaluation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 36 comparisons with regional glacier fluctuations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 37 south of kangersuneq to frederikshåb isblink . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 37 frederikshåb isblink . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 38 nakkaasorsuaq . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 43 sermeq in sermilik icefjord . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 43 isortuarsuup sermia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 45 1cg14004 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 48 kangaasarsuup sermia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 48 north of kangersuneq to saqqap sermersua . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 49 the ice-sheet margin at isukasia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 51 saqqap sermersua . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 51 discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 53 temperature changes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 53 older neoglacial (6500–2000 years bp) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 53 younger neoglacial (2000 years bp – present) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 55 concluding remarks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57 acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 59 references . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 61 index to place names . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 65 5 abstract weidick, a., bennike, o., citterio, m. & nørgaard-pedersen, n. 2012: neoglacial and historical glacier changes around kangersuneq fjord in southern west greenland. geological survey of denmark and greenland bulletin 27, 68 pp. the nuup kangerlua region in southern west greenland became deglaciated in the early holocene and by the mid-holocene, the margin of the inland ice was located east of its present position. discussion of late holocene changes in the frontal positions of outlets relies on descriptions, paintings, photographs, maps, data from investigations of norse ruins, aerial photographs and satellite images. the kangiata nunaata sermia glacier system has receded over 20 km during the last two centuries, indicating a marked response to climatic fluctuations during and since the little ice age (lia). a large advance between 1700 and 1800 was followed by rapid recession in the first half of the 1800s. limited data from c. 1850–1920 indicate that although the long-term position of the glacier front remained c. 10–12 km behind the lia maximum, the late 1800s and the early 1900s may have seen a recession followed by an advance that resulted in a pronounced moraine system. the ice-dammed lake isvand formed during the lia maximum when meltwater from the western side of kangiata nunaata sermia drained to the ameralla fjord in the west. this is in contrast to the drainage pattern before the 1700s, when water probably drained to kangersuneq in the north. thinning of kangiata nunaata sermia resulted in total drainage of isvand between 2000 and 2010 and the discharge of water through austmannadalen has now returned to the same level as that in medieval times. other outlets in the region, such as akullersuup sermia and qamanaarsuup sermia have varied in phase with kangiata nunaata sermia, but with amplitudes of only a few kilometres. in contrast, narsap sermia has been nearly stationary and kangilinnguata sermia may have advanced until the middle of the 1900s. lowland marine outlets in south-western greenland were characterised by large amplitude changes during the neoglacial. extreme examples, in addition to kangiata nunaata sermia, are eqalorutsit killiit sermiat at the head of nordre sermilik fjord in southern greenland and jakobshavn isbræ in disko bugt, central west greenland. the neoglacial advances appear to have occurred at different times, although this may in part reflect the limited information about fluctuations prior to the 1930s. the differences could also reflect variations in mass balance of different sectors of the ice sheet, different subglacial dynamics or topographical factors. the lowland areas are separated by uplands and highlands that extend below the marginal part of the inland ice; in such areas, the outlets have been advancing almost up to the present, so that the position of the glacier front around ad 2000 broadly coincides with the lia maximum. charting the fluctuations of the outlets thus illustrates the large variability of the glaciers’ response to changing climate but it is notable that the number of advancing outlets has decreased markedly in recent years. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: awe@geus.dk mailto:awe%40geus.dk 66 introduction in recent years, the greenland ice sheet (the inland ice) has become a symbol of climate change. the ice sheet is losing mass at an accelerating rate (dahl-jensen et al. 2009; rignot et al. 2011; van as et al. 2011), and it is becoming increasingly clear that we need a better understanding of the past behaviour of the ice sheet and in particular its sensitivity to climatic change. here we present a review of our current knowledge about late holocene changes of the inland ice margin in the nuup kangerlua region in southern west greenland (figs 1, 2). recent climate change has resulted in thinning of marginal parts of the inland ice (dahl-jensen et al. 2009). politicians as well as ordinary citizens are concerned that the melting of the inland ice and the consequent rise in sea level will lead to world-wide environmental and sociological problems. over the next century, the sea level may rise by about 1 m; a large part of this predicted rise is attributed to melting of the inland ice. however, during the past centuries different sectors of the margin of the inland ice have responded to climate change in various ways. some sectors have been stable, others have receded and others have advanced. this implies that predicting glacier hazards such as changes in calf-ice production from the large tide-water outlets or ice-margin changes at hydropower plants can be difficult. it must be remembered that the inland ice is an immense body of ice. at the margin, the ice sheet is often confluent with minor ice caps or other local glaciers that have their own mass budget. the inland ice has extensive accumulation areas, and it is difficult to delimit the catchment areas for the individual sectors or outlets from the ice sheet. in addition, the catchment areas may change over time. the aim of this work is to compile information about marginal changes of the inland ice in the nuup kangerlua region to give an impression of former variations. the work focuses on the last centuries, but includes data concerning the last 6000 years. nuup kangerlua (godthåbsfjord) is the largest fjord complex in southern west greenland and the longest fjord extends c. 160 km. the fjord branches are often surrounded by steep mountains, but lowlands are also represented. the inner parts of nuup kangerlua were colonised by norse people about 1000 years ago. the norse established the western settlement (vesterbygd) with c. 95 farms and two churches, but they abandoned the region after 350–400 years (arneborg 2004; dugmore et al. 2012). the only written information about the greenland ice sheet from the norse period is found in the king’s mirror from c. 1260 where it is briefly stated that most of the land is icebound (larson 1917; weidick & bennike 2007). some legends from the later eskimo people (thule culture) refer to their relationships with the norse people (e.g. accounts by aron from kangeq and jens rosing presented in birket-smith 1961), but provide no information on glacier extent or glacier changes. the european exploration of the baffin bay and davis strait fig. 1. map of greenland showing the location of the studied region and the location of selected ice cores. i n l a n d i c e canada northgrip nuuk summit baffin bay ice core town main area study area comparison for camp century dye-3 kangerlussuaq airport disko sisimiut narsarsuaq 500 km 40°w48°w 72°n72°n 66°n 60°n 32°w 66°n 78°n 8°w40°w56°w 24°w72°w 78°n fig. 1 7 region from c. 1500–1700 did not increase our knowledge of the greenland ice sheet or glaciers in greenland. the region was again colonised by europeans in 1721, and from this time onward a fairly rich body of historical information exists about the outlet glaciers from the inland ice in the region. this presentation summarises information about glacier changes from geological observations, archaeological evidence, historical sources, as well as more recent evidence from aerial photographs and satellite images. an index of relevant place names is included to aid understanding of older data sources. the emphasis is on the scattered historical information on glaciers, which dates back to the early part of the exploration history in the beginning of the 1700s, and up to the first half of the 1900s. work on the fluctuations of the glaciers in this area was initiated 52 years ago (weidick 1959); more recent work and the recognition or re-evaluation of historical sources have necessitated an updating and revision of the record. since the advent of aerial photography in the 1930s and particularly with the recent influx of satellite information, data on ice-sheet history have become both more accurate and more readily available. these data are dealt with in more general terms, tracing the major trends of the glacier fluctuations up to the first decade of this century. setting this study is centred on the area at the head of nuup kangerlua (godthåbsfjord), c. 80 km east of the town of nuuk (fig. 3). the ice-free land in this part of greenland is 100–125 km wide, with lowlands and uplands with elevations up to 1000–1500 m above sea level (a.s.l.). the landscape is dissected by numerous straits and fjords, of which the longest, nuup kangerlua, reaches from the outer coast to the inland ice. in its inner part, in the kangersuneq area, the ice-sheet margin reaches sea level, or close to sea level, in five outlets. from south to north, these are kangiata nunaata sermia, akullersuup sermia, qamanaarsuup sermia, narsap sermia and kangilinnguata sermia (figs 2, 3). fig. 2. map of south-western greenland showing outlet glaciers, the marginal areas of the inland ice and major fjords mentioned in the text. 68°n 70°n 66°n 64°n 62°n 60°n 70°n 68°n 66°n 64°n 200 km 50°w 45°w 50°w55°w 72°n in nuuk kangiata nunaata sermia qamanaarsuup sermia kangilinnguata sermia kangersuneq narsap sermia kangaasarsuup sermia akullersuup sermia frederikshåb isblink saqqap sermersua jakobshavn isbræ nuussuaq disko nakkaasorsuaq qassimiut lobe torsukattak ikerasaap sullua disko bugt isortuarsuup sermia sermeq sermilik buksefjorden allumersat kangerlussuaq airport arsuk bræ nordre qipisarqo bræ sermeq eqalorutsit killiit sermiat qajuuttap sermia sermilik isukasia inland ice nordre isunnguata sermia narsarsuaqnarsarsuaq søndre isortoq paakitssup ilorlia qalerallit sermia søndre sermilik isvand 1cg14004 sisimiut 88 langva kangerluarsunguup illorsuit (13) akul lers uaq nuna taa rsu k umiivik (15) isortuarsuk kangaarsarsuk (55) nipaatsoq (54) austmannadalen gytjesø nikku johannes iversen sø tummeralik (37) kilaarsarfik (51) nansens teltplads (52) tummeralip tasersua kapisillit karra nunatarsuaq saqqannguaq (11) puilasoq 10 km 51°w 51°w 65°n 64°n 50°w 65°n 64°n 50°w kangilinnguata sermia kangersuneq uj ara ssu it p aa va t karra austmannadalen kapisillit narsap sermia aku ller suu p s erm ia kangiata nunaata sermia qamanaarsuup sermia kangaasarsuup sermia 1cg14004 ameralla kangerluarsunguup nuup kang erlu a kapisillit kangerluat ameralik tasersua nunatarsuaq akul lers uaq nuna taa rsu k isortuarsuk itill eq tasersuaq gytjesø nikku isvand johannes iversen sø langvand tummeralik (37) kilaarsarfik (51) saqqarsuaq (16) umiivik (15) ujarassuit (7) puilasoq (8) nansens teltplads (52) tummerallip tasersua illorsuit (13) narsaq (12) saqqannguaq (11) kangaarsarsuk (55) nipaatsoq (54) norse ruin group church ruin amitsua rsuk fig. 3 9 kangiata nunaata sermia and akullersuup sermia are both calving tidewater glaciers with appreciable calf-ice production. the flux of kangiata nunaata sermia was c. 6 km3 per year in 1996, but by 2005 it had increased by 33% (rignot & kanagaratnam 2006). according to mortensen et al. (2011), the glacial ice discharge is 8 km3 per year. at their maximum extent, the two glaciers kangiata nunaata sermia and akullersuup sermia were coalescent, and the calving front of the glacier system was located more than 20 km farther to the north than today. the term ‘kangiata nunaata sermia glacier system’ is used here to describe the two confluent glaciers. the third glacier, qamanaarsuup sermia ends on land today, but previously ended in a proglacial lake that was dammed by the advanced, confluent glacier system consisting of the present-day kangiata nunaata sermia and akullersuup sermia. the fourth glacier, narsap sermia, is also a tidewater glacier, but calf-ice production is small, and it is sometimes possible to sail along the glacier front by boat. the fifth outlet glacier, kangilinnguata sermia, is situated at the head of ujarassuit paavat, which is a northern branch of kangersuneq fjord at the head of nuup kangerlua. kangilinnguata sermia is presently a minor land-based outlet. whereas kangiata nunaata sermia, akullersuup sermia and qamanaarsup sermia were surrounded by broad, fresh trimline zones in the last century, such a zone is only found to a very restricted degree around narsap sermia. the trimline zone around kangilinnguata sermia is narrow, which indicates a slight recession of the glacier in recent times. kangaasarsuup sermia, south of kangiata nunaata sermia, has experienced a somewhat larger glacier recession (fig. 4). mass-balance measurements have been carried out on qamanaarsup sermia between 1979 and 1989 and on a small local nameless glacier (1cg14033) situated at the junction between ameralik fjord and kangerluarsunnguaq (buksefjorden), c. 75 km south-west of qamanaarsuup sermia, during the period 1982–1989 (braithwaite 1989, 1990; braithwaite & olesen 1989). at both places, the investigations were carried out to evaluate the hydropower potential in the region. on qamanaarsuup sermia, volume changes were also measured from 1968 to 1980 (knudsen 1983) and again from 1988 to 1993 (taurisano 2004). north of the kangersuneq area (at isukasia), glaciological investigations were conducted in connection with evaluations of the hydropower potential at the inland ice margin in the years 1974–1980 by the arctic consultant group (acg) and vattenbyggnadsbyrån (colbeck 1974; kryolitselskabet øresund 1980). notes on the holocene history of the area the holocene history of the area is not known in detail. a quaternary map of the region on a scale of 1:500 000 was published by the geological survey of greenland (beschel et al. 1978), but without explanatory notes or a map description. notes on the quaternary history have been published by weidick (1975a, b) and long et al. (2006). during the last glacial maximum, c. 21 cal. ka bp (cal. ka bp = calibrated to calendar ka before present, where present = ad 1950), the margin of the inland ice extended to the edge of the continental shelf according to vinther et al. (2009). the outer coast of the region became ice free during the early holocene (bennike & björck 2002). during the following millennia, net recession continued, but ice-margin deposits mark some halts or minor readvances of the ice margin. the correlation of these ice-margin deposits in western greenland is uncertain (weidick 1968, 1984a, b; kelly 1980, 1985; long et al. 2006). the recession of the ice margin over the present ice-free land followed after the abrupt warming at the younger dryas – holocene transition, which is dated to 11.7 ka before ad 2000 (rasmussen et al. 2006), and which was followed by increasing temperatures during the earliest holocene. a characteristic feature of the kangersuneq area is the wide belt of ice-margin deposits, which can be followed at the heads of nuup kangerlua and ameralik fjords at a distance of 10–30 km outside the present margin of the ice sheet. the deposits indicate a halt or readvance. these holocene ice-margin features have been variously termed the fjord stage by weidick (1968), the younger fjord moraine system by ten brink & weidick (1974) and the kapisigdlit moraine system by kelly (1985). the term ‘the kapisigdlit stade’ is used here to stress the local character of these ice-margin deposits. near the settlement of kapisillit, the marine limit drops abruptly from facing page: fig. 3. map of the inner nuup kangerlua and kangersuneq fjord region, showing the locations of named features and norse ruins mentioned in the text. the map is based on aerial photographs from 1985; outlined area shown in fig. 16. 1010 gytjesø n aar suk kangerluarsunguup isortuarsuk kapisigdlit stade austmannadalen kapisillit nunatarsuaq akul lers uaq ’lake 8 m’ nikku karra tasersuaq 8.2 0.2 9.8 8.5 10.8 51°w 51°w 65°n 64°n 50°w 65°n 64°n 50°w kangilinnguata sermia kangersuneq uj ara ssu it p aa va t karra austmannadalen kapisillit narsap sermia akull ersu up ser mia kangiata nunaata sermia qamanaarsuup sermia kangaasarsuup sermia 1cg14004 ameralla kangerluarsunguup nuup kang erlu a kapisillit kangerluat ameralik tasersua nunatarsuaq akul lers uaq nuna taa rsu k shell date basal gyttja from isolation basin church ruin (kilaarsarfik) isortuarsuk 6.4 5.2 5.0 4.3 7.6 kapisigdlit stade itill eq tasersuaq kapisigdlit stade covered during the little ice age gytjesø ’lake 8 m’ nikku isvand fig. 4 4.9 amitsua rsuk 10 km 11 over 100 m a.s.l. west of the moraine system to less than 80 m a.s.l. east of the moraines. on the basis of the radiocarbon ages from the area (fig. 4) and archaeological information on increasing relative sea level at the church ruin at kilaarsarfik (‘sandnes’) after medieval times (fig. 3; ruin group no. 51; roussell 1936), an uplift curve has been proposed (fig. 5). in the disko bugt region, the correlative of this c. 9–8 ka old stade was called the ‘fjord stage’ by weidick & bennike (2007, p. 37). according to the uplift curve, the kapisigdlit stade can be dated to 8.1–8.3 ka bp, and it may reflect a readvance as a result of the 8.2 ka bp cold event (alley et al. 1997). the age of the kapisigdlit stade may correspond to the age of the tasiussaq moraine system in disko bugt (weidick & bennike 2007, young et al. 2011). however, a shell of the marine bivalve macoma calcarea, found at kangersuneq near the settlement of kapisillit (figs 4, 6) was dated to 9490 ± 105 14c years bp (ua-3476; fig. 6b), or 10.6–11.1 cal. ka bp (table 1). the shell was found in a 5–10 m high coastal cliff section with a surface that gradually rises to 13 m a.s.l. at the foot of the proximal side of a kapisigdlit stade moraine, in a silty boulder diamicton containing concretions with shells (fig. 6b). the shell fauna included rare but well-preserved shells of mya truncata and balanus sp. the shells probably come from invertebrates that lived at the bottom of kangersuneq icefjord at a time when the ice margin was located near its present position or farther to the south and east than today. during the kapisigdlit stade, the shells were dredged from the fjord bottom by the advancing glacier, and deposited in a diamict accumulation, which is now exposed in a coastal cliff section. moraines related to the kapisigdlit stade were first described from nansens teltplads (nansenip tupeqafia) at the entrance to austmannadalen from ameralla (near norse ruin group 52c; figs 3, 7). they were called facing page: fig. 4. map of the inner nuup kangerlua and kangersuneq fjord region, showing locations of radiocarbon-dated samples (ages in calibrated ka bp, tables 1, 2), the approximate position of the kapisigdlit stade (weidick 1975a) and the maximum extent of the ice margin during the little ice age (lia max). the red areas around the present ice margins show the extent of the trimline zone. this zone is broadest around the kangiata nunaata sermia glacier system. the map is based on aerial photographs from 1985. 20 40 60 80 100 el ev at io n (m a .s. l.) age (cal. years bp) 40006000800010 000 2000 0 0 marine shells archaeological dating basal gyttja from isolation basin fig. 5 fig. 5. model of holocene relative sea-level changes in the kangersuneq icefjord area, based on data presented in table 1; a.s.l.: above sea level. locality lat. long. elevation laboratory age calibrated age material reference (north) (west) m no. 14c yrs bp * ±1 stdv., bp † * ages were not normalised for isotopic fractionation, except for ua-3476 that was normalised to a δ13c value of 0‰ on the pdb scale. † calibrated according to the calib09 dataset (gyttja) and the marine09 dataset (marine shells); stdv: standard deviation. johs. iversen sø 64°24´ 50°12´ 100 k-2294 8640 ± 130 gyttja fredskild 1983 gytjesø 64°23.5´ 50°21.5´ 57.3 k-2295 7430 ± 100 gyttja fredskild 1983 kapisillit 64°26´ 50°10´ 40 k-1036 7560 ± 150 shells weidick 1968, 1972 amitsuarsuk 64°32´ 50°28´ 16 i-8596 6670 ± 110 shells weidick 1976 kapisillit 64°28´ 50°11´ 5 ua-3476 9490 ± 105 macoma calcarea this study lake 8 m 64°26´ 50°12´ 8 k-802 4340 ± 120 gyttja fredskild 1973 kilaarsarfik 64°15´ 50°13´ –2 archaeol. roussell 1936, 1941 table 1. selected radiocarbon age determinations from the kangersuneq region 9432–10153 8031–8402 8106–8861 7365–7776 10570–11109 4583–5305 c. 800 1212 ‘jøkelgjærde’ (moraine ridge) by nansen (1890, map) and bruun (1917, p. 102). this moraine may be cut by a marine terrace at c. 50 m a.s.l. another early observation of icemargin deposits come from kangersuneq fjord, near ruin group 15 (umiivik), where roussell mentioned that about 700 m up from the ruin site is a large moraine bank running across the valley (figs 3, 7; roussell 1941, pp. 60–61). the recession presumably continued after the kapisigdlit stade, up to the peak of the holocene thermal maximum at c. 7–6.5 ka. at around this time the ice margin was generally located inland of the present position. it is not known how far east the ice margin retreated. however, a minimum estimate can be given for the time interval during which the ice margin was behind the present margin, based on ages of marine shells in concretions sampled from little ice age moraines and from the alluvial plains in front of the outlet of kangilinnguata sermia, and dredged by the glacier from the bottom of the fjord. age determinations of four samples gave ages of 6.4 to 4.2 cal. ka bp (table 2). it is presumed that the net readvance was characterised by separate readvances of increasing magnitude, which generally culminated towards the end of the little ice age at ad 1700–1900. the geomorphology of moraine localities and the content of fossils in the deposits were described by the german geologist k. gripp (1932, 1975), who also described marine fossils from alanngorlia at sermeq glacier and frederikshåb isblink. 100 km disko 63°n 60°n 52°w 48°w 66°n 56°wa lgm s 11 ka 8–9 ka lowland highland kapisillit 69°n b nuuk 1 km 64°25´n 50°10´w johannes iversen sø ua-3476 (10.8 ka) kapisillit kapisillit kangerluat kangersuneq kapisigdlit stade 64°25´n 50°10´w fig. 6 1968 fig. 6. a: ice-margin positions during the recession of the of the inland ice margin in western greenland (simplified). the numbers indicate approximate mean ages in ka bp. lgm: last glacial maximum, dated to c. 22 ka bp, the position is according to vinther et al. (2009). s: the sisimiut glacial event. the latter stade was dated to 11–14 radiocarbon ka bp, corresponding to c. 12.9–17 cal. ka bp according to kelly (1985). however, it is possible that the sisimiut glacial event may correlate with the younger dryas cooling (11.7–12.9 cal. ka bp; rasmussen et al. 2006). the stade marked 8–9 ka shows the approximate position of the ‘fjord stade’, including the local kapisigdlit stade in the kangersuneq – nuup kangerlua area and the marrait and tasiussaq moraine systems in disko bugt. lowlands are areas below 300 m a.s.l. b: aerial photograph showing the area around the isthmus at kapisillit between the fjords kangersuneq and kapisillit kangerluat. ice-margin features belonging to the kapisigdlit stade, the location of the shell sample dated to 10.6–11.1 cal. ka bp (ua-3476) and johannes iversen sø, investigated by iversen (1953) and fredskild (1973, 1983) are shown. aerial photograph from 13 august 1968, geodetic institute, route 281r, no. 146. 13 ruin group 16 fig. 7 fig. 7. norse and eskimo ruin groups in the area between nuup kangerlua (godthåbsfjord) and ameralik fjord (bruun 1917); bruun’s numbers for ruin groups are still being used today. d. bruun used a map by j.a.d. jensen (drawn after field work in 1884 and 1885) as a basis for plotting archaeological sites after collecting information in greenland in 1903 and using additional information from o. bendixen after his travels in 1916. the map shows that the norse farms and churches are found in the interior part of the ice-free land. jensen’s map was in general use up to the 1930s. legend: 1. norse ruins. 2. ruins previously believed to be norse. 3. eskimo tent sites. 4. eskimo winter houses. 5. eskimo winter houses from the time of hans egede (younger than 1721). 6. inhabited sites (i.e. from 1884 to the 1930s). the red dot marks the location of ruin group 16 according to roussell (1941). 1414 moraine systems (partly with ice-cored moraines) are especially well developed around the kangiata nunaata sermia glacier system. relicts of shear moraines can be seen in kangersuneq c. 11–12 km behind the little ice age maximum extent as defined by the extension of the trimline zone. in the following, this advance is called the ‘1920 stade’. the first overview of the bathymetry of the nuup kangerlua region was compiled by beschel et al. (1978), although data were lacking from major parts of the interior fjords. mortensen et al. (2011) presented a depth profile of nuup kangerlua and kangersuneq. the depth of the inner part of nuup kangerlua is about 600 m, whereas the depth in kangersuneq icefjord is around 300–400 m (fig. 8). a major threshold is found at the little ice age maximum, where the depth is c. 150 m. another smaller threshold is found c. 8 km inside this, with a depth of c. 300 m. the latter is found 1–2 km in front of the 1920 stade. a third threshold may occur at the 2010 frontal position of kangiata nunaata sermia, where the water depth is 210–240 m. the thresholds are separated by depressions with depths of 350 and 400 m (fig. 8). the depressions may have been eroded during time periods when qamanaarsuup sermia (during the kapisigdlit stade) and akullersuup sermia (during the little ice age maximum) were tributaries to the kangiata nunaata sermia glacier system. the present frontal height of kangiata nunaata sermia is <50 m; the glacier front is probably mainly resting on the fjord bottom. geological, archaeological and historical information on glacier fluctuations and sea-level changes the uplift curve (fig. 5) shows a rapid emergence of land in the early holocene, which was a consequence of the decrease of the ice load over the region. uplift was followed by submergence in the late holocene. indications of relative submergence were already noted by arctander (1793) and pingel (1841, 1845) based on observations of archaeological sites situated at or below the present sea level. their observations were confirmed by matthiassen (in: gabel-jørgensen & egedal 1940, pp. 8–10), roussell (1936, 1941) and larsen & meldgaard (1958) based on laboratory age calibrated age reference 14c yrs bp * ±1 stdv., bp † ua-3473 5590 ± 90 6240–6628 this study ua-1089 4535 ± 110 4951–5544 weidick 1993 ua-3475 4280 ± 95 4683–5244 this study ua-3474 3780 ± 70 4057–4448 this study table 2. neoglacial radiocarbon ages of shells from the bottom of ujarassuit paavat (64°50´n, 50°0´w) * the radiocarbon ages were normalised for isotopic fractionation to a δ13c value of –0‰ on the pdb scale. †calibrated according to the marine09 dataset; stdv: standard deviation. no. 200 –200 –400 400 600 800 kapisigd lit st ade present sea level sea level 8.2 ka m 1960s 1920 stad esaqqarsuaq ruin group 16 lia maximum qamanaarsuup tributary sermia akullersuup tributary sermia 0 5 10 15 20 km lia maximum fig. 8fig. 8. depth profile of kangersuneq icefjord showing profiles of kangiata nunaata sermia during the kapisigdlit stade (8.2 cal. ka bp?), during the little ice age (lia) maximum (1700s), during the 1920 stade (c. 1920) and during the 1960s. the minor depression of the fjord bottom 4–6 km off qamanaarsuup sermia and c. 14 km off akullersuup sermia may be due to the erosive effect of the frontal part of kangiata nunaata sermia glacier system when the two tributary outlets joined the main glacier kangiata nunaata sermia. 15 archaeological investigations, as well as by saxov (1958, 1961) based on geophysical work. a review of these investigations was published by weidick (1996). investigations have been carried out at a few localities in southern west greenland by dating the timing of isolation of lakes from the sea using pollen analysis and radiocarbon age determination. the pioneers of such investigations in the area were iversen (1953) and fredskild (1973, 1983). more recent investigations have been made by long et al. (2009) at a locality on the outer coast south of sisimiut, and by bennike et al. (2011) at a locality near sisimiut. the relative sea-level history in the inner part of the nuuk region was undoubtedly different from that of the sisimiut region 300 km to the north-west, but the fact that the church ruin at kilaarsarfik (‘sandnes’) is now covered by the sea at high tide shows that this region has been transgressed. dietrich et al. (2005) suggested a recent uplift of about 1 mm per year near sisimiut and a subsidence of 2.2 mm per year at kapisillit in the inner region of nuup kangerlua based on gps measurements at a number of stations in western greenland. long et al. (2009) suggested that the difference reported by dietrich et al. (2005) could be a consequence of the behaviour of the ice margin in the two areas. more recently, bevis et al. (2012) reported an uplift rate of 7.5 mm per year at kapisillit. although this figure may be uncertain due to problems with uplift processing, it is likely that the uplift rate has accelerated. roussell (1936, p. 8–10; 1941, p. 14–18) in his treatise on the norse farms and churches in the inner parts of nuup kangerlua and ameralik made two important observations on the environmental situation of the norse settlements: 1. the relative sea level was some metres lower at the kilaarsarfik church (ruin group 51, figs 3, 7) at the head of ameralik fjord in early medieval time than today. we suggest that the following relative sinking of the land is primarily related to the increasing glacier load during the neoglacial. 2. the many farms along the head of ameralik, in austmannadalen west of kangiata nunaata sermia, at the head of ujarassuit paavat north of kangersuneq and in kangersuneq icefjord must have been more accessible from the sea in medieval times than at present, where kilometre-long tidal flats or calf-ice accumulations are found at the heads of the fjords. the tidal flats reflect a large increase in sediment supply, probably due to a marked glacier advance after the norse disappeared from the area, presumably around 1345 when the norwegian priest ivar bar darsson visited the area (gad 1967; arneborg 2004). as mentioned in the introduction, data on glaciers and glacier changes are sparse up to the first half of the 1800s. however, in the mid-1800s there was a growing international interest in the arctic regions. the first systematic investigations of western greenland were undertaken between the 1870s and 1900 by the commission for the direction of the geological and geographical investigations in greenland. in addition to geological and geographical studies, archaeological, botanical and other scientific investigations were carried out. the results were published in meddelelser om grønland, which was issued by the commission. data from aerial photographs and satellite images the use of aerial photography in greenland began in the 1930s, in connection with a new systematic mapping programme (wenzel-petersen 1970; nielsen 2000). the oldest photographs were oblique ones. during the second world war, the us air force covered parts of greenland with a combination of oblique and vertical photographs (trimetrogon). oblique aerial photographs were also used by the geodetic institute in following years. these photographs cover large areas, but distant objects show little detail and more recent series are thus all vertical. aerial photographs from greenland have been acquired and collected since the 1930s by the former geodetic institute which was established in 1928 and was originally an institution under the war ministry (krigsministeriet; helk 1954; wenzel-petersen 1970, p. 31). the geodetic institute produced and sold maps and copies of aerial photographs. the total collection of aerial photographs was available without restrictions, even during the cold war. on 1 january 1989, the geodetic institute became part of the national survey and cadastre (kort & matrikelstyrelsen), currently in the ministry of the environment (nielsen 2000, p. 77). in addition to topographical mapping, the west greenland photographs from the 1930s were also used in connection with archaeological investigations of the norse ruins. the many details seen on the photographs made them suitable for archaeological and other scien1616 tific investigations. however, the old series did not cover all the ice-free land, and it was not until 1985 that a total coverage of vertical aerial photographs was achieved. since the 1970s, satellite images have been used to map glacier changes. these images cover large regions and are available at short intervals. data from satellites can be used in mass-balance calculations of specific sectors (dahl-jensen et al. 2009). however, in this work we have only used data from satellites to update information from historical sources and aerial photographs. fig. 9. a: isvand and the eastern part of austmannadalen in 1888. eastern part of a map by o.c. dietrichson (nansen 1890). skridjökel: glacier, here part of a southern outlet of the inland ice (kangaasarsuup sermia). isvand is drawn fairly schematically, but at its east side one can see the damming glacier marked by icebergs in the lake. b: isvand and its discharge into the river through austmannadalen. vertical aerial photograph from 16 august 1968, geodetic institute, route 281t, no. 85. the arrow indicates the island discussed in the text. the recession of the damming ice margin is given by approximate positions of the ice margins with years of observations from 1888 to 1985. the recession was probably interrupted by minor advances, with the most pronounced culminating just before 1920 (marked 1920 stade). the most impressive change occurs in the area between the present glacier margin and the limit of the maximum ice coverage during the little ice age. this limit is marked lia max, and the area behind is still relatively barren even though major parts of this trimline zone have been free of ice for around 200 years. the position of the front of kangiata nunaata sermia was stable between 1985 and 2010. after weidick & citterio (2011). 1955 1948 1948 1936 1985 1968 1936 1888 lia max 19 20 st ade lia max 1968 1 kmka ng aa sa rs uu p se rm ia kangiata nunaata sermia austmannadalen isvand langvand 64°14´n 49°40´w 64°14´n 5 km 64°14´n 49°35´wa b fig. 9 1968 17 glacier changes in the kangersuneq area the description and evaluation of glacier changes primarily concern the fronts of the glaciers kangiata nunaata sermia, akullersuup sermia, qamanaarsuup sermia and kangilinnguata sermia although the record from kangiata nunaata sermia is supplemented by data from the locality of isvand, a former ice-dammed lake situated on the western flank of the glacier (figs 3, 9). the individual glaciers or localities are treated separately, the presentation of the historical and modern data being followed by an evaluation of these datasets in terms of the glacial history. the former ice-dammed lake isvand the description below concentrates on the area east of austmannadalen, in particular the area including the former ice-dammed lakes isvand and langvand, located south-west of kangiata nunaata sermia and north of kangaasarsuup sermia (figs 3, 9). the earliest information about the ice margin in this area was given by hans egede (1925, p. 104) in 1723. he refers to observations by reindeer hunters, who from ameralik fjord presumably travelled up austmannadalen, and observed the large ‘ice mountain’ which expanded over the mountains and undoubtedly was connected with the ice cover and mountains that egede had seen farther south. the hunters saw a large lake under the ‘ice mountain’, into which large ice lumps had fallen from the ‘ice mountain’, and egede compared it with an earlier mentioned ‘ujarachsuach’ (now: kangersuneq), which also was full of icebergs that had fallen from the ice. it is uncertain whether this description refers to isvand, but isvand was the largest of the ice-dammed lakes in this area at the head of austmannadalen. this ice-dammed lake formed after the little ice age maximum extent of the ice, when the whole area of the later isvand was ice covered. the comparison to the ice conditions in ‘ujarachsuach’ fjord is also open to question. it appears that this name was sometimes used for the present kangersuneq, but usually it was only used for the fjord ujarassuit paavat (figs 3, 10). it may also have been via austmannadalen that governor claus enevold paars in april 1729 visited the ice margin, sailing from nuuk to the head of ameralik and then walking through austmannadalen with a party of danes, guided by two greenlanders to reach the inland ice. from his observations near the ice margin, he described glacier crevasses, the crystalline structure of the glacier ice, and erratic boulders on the ice surface (bobé 1936a, p. 26, 1936b, pp. 186–189). on the trip he also found ruins, but he does not mention the ice-dammed lakes at all. if large ice-dammed lakes existed, they must have formed an obstacle for approaching the inland ice margin and would probably have been recorded. egil thorhallesen, a missionary who lived in greenland between 1765 and 1775, toured the central parts of western greenland, especially in 1774 and 1775. he published an account of the norse ruins (thorhallesen 1776), but his original description of the country, its people and nature was not published until 1914 by bobé (thorhallesen 1914, including maps of the area). thorhallesen mentioned several norse ruins, amongst others at a locality called kangia, presumably at the head of ameralla: in the eastern end of the bay several ruins from the old europeans are also found; some may also be situated on high ground and some may have been covered by the glacier-ice in recent time. this is described in more detail in the “efterretning om rudera i baals revier [treatise on rudera in baals revier]”. (thorhallesen 1914, pp. 51–52; authors’ translation.) it is difficult to locate thorhallesen’s place names from his map (fig. 10). ‘baals revier’ is nuup kangerlua (godthåbsfjord). the river named ‘laxelv’ is probably the river running through austmannadalen. ‘storelv’ is probably the naajat kuuat river, which drains the glacier kangaasarsuup sermia (not shown on thorhallesen’s map, fig. 10) and also drains the large lake isortuarsuk. the head of ameralik and ameralla fjords was called ‘kangia’ by thorhallesen, but in a note by bobé (thorhallesen 1914, p. 51) it was changed to ‘umiviarsuit’, which is a locality close to the kilaarsarfik church ruin (bruun’s map 1917, fig. 7). thorhallesen noted that the length of ameralla fjord from eqaluit to its head at kangia (or umiviarssuit) is 2 danish miles (15 km), which is close to the distance on modern maps. on thorhallesen’s map, the lengths of the fjords are generally close to distances on modern maps whereas the orientations of the individual parts are distorted. roussell (1941) noticed that the route to the norse ruins behind the present extensive tidal flats may have been easier in the 18th centu1818 1 2 5 3 4 6 7 8 9 10 fig. 10 19 ry than today. it is possible that much of the sedimentation took place after thorhallesen’s visit. sedimentation rates may have increased due to the increased ablation that accompanied glacier recession in the 1800s. the ruins at tummeralik were also mentioned by the early explorers of the region, and from a visit to the area around tummeralik on 1 june, 1810 (giesecke 1910, 254–259). giesecke described the ruins at tummeralik as well as the ice margin of the ‘eisblink’, which may be the ice margin situated c. 5 km east of tummeralik (tummerallip tasersua). ruin group 37a shown on bruun’s map (fig. 7) is not mentioned or plotted on newer maps of norse ruins in the area (gad 1967; kort & matrikelstyrelsen 1993) and may therefore not exist, at least not as norse ruins. from a camp at ‘auaitsirksarbik’ (3–4 km from the ice margin of the glacier?) in 1810, giesecke (1910, pp. 257–258) described the ice margin: this evening i went to the incredibly beautiful glacier (sermersoak in greenlandic), which is surrounded by a lake. the glacier is up to 80 feet high. its margin is split into pointed cone-shaped prisms that are separated by deep crevasses, and some of the cones are almost pyramid-shaped. the surface of the glacier has large, often undulating peaks and depressions. the first impression is ghostlike; and the surrounding, completely barren area put me in a melancholy mood; it seemed to me that i was in another world. it was midnight – i was resting alone in this depressing area, which had presumably never before been visited by any european. however, soon i enjoyed an uplifting drama in this desert when the sun rose behind this colossal mirror of ice... (authors’ translation). further observations by giesecke may refer to the development of a trimline zone and icemarginal lakes, but his description is unclear. several old maps of the area, drawn by local greenlandic hunters, are housed in a collection at the royal library in copenhagen. this collection was established on the initiative of hinrich rink, the inspector and later director of the royal greenland trade. he encouraged local hunters to collect information and draw maps of their hunting areas. the linguist and missionary, samuel kleinschmidt, also participated in this collection of information; his map of the region (fig. 11) provides the best overview of the area from this time. the map was presumably drawn in the years before 1859, based on various sources and printed in 1860 (kleinschmidt 1860). on this map, a lake is apparently indicated at the site where austmannadalen ends at the glacier margin, although it is possible that this lake is just two converging rivers. on newer versions of the map from the last half of the 1800s, this feature was sometimes coloured to indicate a lake. kleinschmidt’s map was based on maps prepared by local hunters and their information, as well as on his own sketches, bearings and data from surveyors. his map is recognised as being of a high standard for the time. a comparison of thorhallesen’s map from 1776 (fig. 10) with kleinschmidt’s map from 1860 (fig. 11) shows major improvements. the importance of kleinschmidt with regard to the geographic knowledge of the nuuk region is emphasised by wilhjelm (2001, pp. 144–152). j.a.d. jensen reported from his visit to the area in 1885 (jensen 1889, p. 88) that he could not enter the interior of kangersuneq due to dense calf ice in the fjord. furthermore he noted: in later years the calf-ice concentration [in kangersuneq] has increased considerably, and the greenlanders are now rarely able to travel there, which was quite common earlier. (authors’ translation.) this could mean that the glacier had receded. however, compared with the earlier maps of the 1800s, j.a.d. jensen’s map from his expedition is more accurate and a later version of this map was used as the base map for bruun’s systematic mapping of the norse ruins in the area ( bruun 1917). the map of the interior parts of kangersuneq (fig. 7) may well have been drawn from the mountain ‘nikok’ (jensen 1889, pp. 88–89). according to jensen’s travel report, this mountain is 3130 feet high (970 m), and he noted that he had a good view of the interior landscape around the southern part of kangersuneq. it must be the same mountain, named nivko, which on the map published by kort & matrikelstyrelsen (1993) is shown with facing page: fig. 10. map of the ‘gothaabs destrict’ compiled by e. thorhallesen in 1776, accompanying the report on his voyage to central west greenland (thorhallesen 1914). the map gives an impression of the geographical knowledge gained by the end of the 1700s of the area of greenland that was best known at that time. the numbers show the location of some of the place names used in this study: 1: nuup kangerlua (godthåbsfjord). 2: ujarassuit. 3: kangersuneq. 4: narsap sermia. 5: isthmus between kapisillit kangerluat (on the map: ‘pisigsarfik fjorden’) and kangersuneq. 6: kangiata nunaata sermia. 7: ameralla. 8: austmannadalen. 9: naajat kuuat. 10: eqaluit. the map published in 1914 measures 32 × 23 cm. e. thorhallesen (1734–1789) was born in iceland and was a priest in greenland (1765–1775). he made a description of west greenland in connection with mapping of norse ruins. 2020 a fig. 11 fig. 11. map of the nuuk area (title in greenlandic). the danish text below the map translates: map of the inner part of godthåb district, with special regard to the former scandinavian settlements. compiled from various observations, as well as from drawings and accounts by greenlanders, especially aron and abraham from kangeq. by s. kleinschmidt, godthaab 1860. printed in the printing office of the inspectorate. (authors’ translation.) the danish headings to the right of the map translate as follows: sunde og fjorde: straits and fjords. indsøer: lakes. fjelde m:m:: mountains etc. forskjellige benævnte egne: various mentioned areas. øer: islands. nyere hustomter: newer house ruins. pladser beboede 1859–60: inhabited settlements 1859–1860. the legend translates as follows: betegner … beboede pladser: designates … inhabited settlements. … grønlandske ruiner: eskimo ruins. … nordiske ruiner: norse ruins. the scale (maalestok) is in danish miles (1 danish mile = 7.5 km); a: akullersuaq. the map is reproduced from gulløv (1983). a printed version of the original map is in the royal library in copenhagen; the map in this version measures 26 × 19 cm. s. kleinschmidt (1814–1886) was a teacher at the teacher’s college in nuuk (godt håb seminarium); he published a famous greenland grammar and added many details to the geographical knowledge of west greenland. 21 an altitude of 924 m a.s.l., and which is located on the south-western side of kangersuneq near the settlement kapisillit. jensen also described ice-dammed lakes along the south side of narsap sermia on the semi-nunatak nunatarsuaq and their draining through narsap sermia and he mentioned the innermost semi-nunatak nunataarsuk. in view of the details provided, the approximate position of the front of kangiata nunaata sermia south of the semi-nunatak akullersuaq may be realistic (figs 3, 4), but the front may also have been copied from kleinschmidt’s or other early maps from the middle of the 1800s. the following information about the area between kangiata nunaata sermia and ameralik fjord in 1888 is taken from f. nansen’s description (nansen 1890). nansen crossed the inland ice from south-east greenland to ameralik fjord in west greenland. he described the descent from the inland ice and compiled a map over parts of the ice margin, the ice-dammed lakes isvand and langvand and the valley (austmannadalen) leading down to ameralla fjord (fig. 9). all three names were given by nansen’s expedition. the area was mapped by o.c. dietrichson; only the western part of it concerning isvand is shown here (fig. 9a). this map from 1888 shows that the marginal lakes are situated inside the trimline zone, which must have formed before this year. if kleinschmidt’s map from 1860 shows a lake at this site, isvand must have existed around or before the middle of the 1800s. initial thinning of the glacier implies that recession from the maximum extent of kangiata nunaata sermia began before the visit by nansen’s expedition. the following recession can be seen on aerial photographs from 1936 to 1985 (fig. 9b). the norse ruins in austmannadalen were located in 1934 and excavated in 1937 (roussell 1941, p. 14). in connection with these activities, the area was covered by aerial photographs in 1936. the photograph in fig. 9b from 1968 shows the eastern part of austmannadalen with the ice-dammed lakes isvand and langvand. the glacial extent in 1936 can be compared with the situation in 1888 (nansen 1890, map; fig. 9a), as well as with an aerial photograph from 1985 and an aster image from 2010 (fig. 12). the photograph from 1968 shows the inland ice margin surrounded by a wide trimline zone. it is difficult to differentiate between the ice-dammed lakes isvand and langvand (fig. 9b) and the surrounding trimline zone as the lakes were filled with the same silt that covers the vegetation-poor trimline zone; this is in marked contrast to other lakes in the area that appear black on the photographs. the extent of the trimline zones indicates that the maximum coverage of the margin of the inland ice occurred before 1888, but also that the ice margin at this locality in 1888 had already receded 2–3 km from the little ice age maximum. on the aerial photographs from 1936 and subsequent years, a series of ice-margin features, which must have formed during an advance before the 1930s, can also be seen. they can be followed down to the former semi-stable front of the kangiata nunaata sermia glacier system situated c. 12 km behind the little ice age maximum. the stade marked by these moraines is the 1920 stade mentioned earlier. on the aerial photographs from 1936, an ice-free headland, peeping out from the ice margin in the lake, is seen. following recession until 1985 (fig. 9b), this headland became an island (see arrow on fig. 9b) and the recession of the ice margin during this period must be c. 2 km. the shape of the isvand lake in the period 1936–85 remained unchanged, indicating a nearly permanent lake level at c. 360 m a.s.l. (kort & matrikelstyrelsen 1993), with no periodic sudden drainage of isvand under or through kangiata nunaata sermia to kangersuneq and with a permanent discharge through austmannadalen. comparison of the 1985 photograph with a landsat image from 2009 shows a dramatic change in the area. satellite images from between 1987 and 2010 indicate that the shrunken isvand lake underwent sudden drainage events via kangiata nunaata sermia in 2004 and 2009. it appears that the continuous thinning of kangiata nunaata sermia resulted in drainage of isvand. the fluvial discharge through austmannadalen to ameralla and ameralik fjords must now be just as modest as it was during the norse period. however, the area formerly covered by isvand is now covered by a layer of silt and the head of ameralla is filled up with extensive tidal-flat silt accumulations; these features testify to the former influence of isvand and its drainage through austmannadalen. photographs taken by dirk van as on 26 august 2010 apparently show a renewed filling up of the ponds of the former isvand. this is also seen on aster images from earlier in 2010 (fig. 12). presumably the lake remnants are now partly emptied by periodic drainage via kangiata nunaata sermia. 2222 evaluation during the time of the norse settlers, the river running through austmannadalen only received water from the restricted hydrological basin of austmannadalen and therefore had a very modest discharge. the following culmination of the little ice age is uncertain, but the ice margin was probably situated at the pass point between austmannadalen and the valley around isvand. h. egede visited the region in 1723 and mentioned an expanding ‘ice mountain’ and water ‘under’ (and around?) the ice (egede 1925, p. 104). in 1729, c.e. paars traversed the area, but he does not mention an ice-dammed lake (bobé 1936b, pp. 186–187). it is therefore possible that isvand did not exist and the ice margin may have been located at the pass point between isvand and austmannadalen (at c. 360 m a.s.l.). during the period when isvand existed in the 1700s, 1800s and the 1900s, the pass between isvand and austmannadalen received meltwater from large parts of the western ablation areas of kangiata nunaata sermia that drained via austmannadalen to ameralla. the lake may have had the same level at c. 360 m a.s.l. throughout the period of its existence, and the development of the lake reflects the recession of the ice margin, which led to a growth of the lake area, but without any essential change of its outline. after the rapid drainage of the fig. 12. advanced spaceborne thermal emission reflection radiometer (aster) satellite image of the area in front of kangiata nunaata sermia and former size of lake isvand showing the situation in 2010. the outline of the former lake isvand in 1968 is shown with white lines (see fig. 9b). the approximate positions of the front of kangiata nunaata sermia are shown for the period c. 1920–2010; the akullersuup sermia ice front remained nearly stationary between 1985 and 2010. lia max: little ice age maximum. after weidick & citterio (2011). 2 km 1920 1985 2009 2010 1968 1920 stade lia max lia m ax 193 6 1946? 197 9 1948 2009 19 85 /86 kangiata nunaata sermia kangaasarsuup sermia akullersuup sermia austmannadalen langvand former size of lake isvand 64°16´n64°16´n 49°40´w 49°40´w 19 68 2010 fig. 12 2010 23 lake through kangiata nunaata sermia in 2004 and 2009, the ice margin became more stable, and the fluvial regime in austmannadalen became similar to that in medieval times. the only change compared with the medieval conditions is the blanket of silt and mud that now covers the trimline zone including the earlier lake floor of isvand. in addition, an extensive tidal flat has formed at the head of ameralla fjord. the glacial events in the austmannadalen and isvand area can be summarised as follows: c. 1700 – c. 1750: the advance leading to the little ice age maximum presumably occurred c. 1729 and led to the formation of isvand, the ice-dammed lake, in perhaps 1723. this implies that significant drainage of meltwater through austmannadalen from kangiata nunaata sermia took place. subsequent minor recessions and readvances of the ice margin may have occurred during this period. c. 1750 – c. 1800: initial thinning of the western flank of kangiata nunaata sermia. the size of isvand depended on the minor fluctuations of the glacier. over this time period, the area drained via austmannadalen to ameralla. the size of isvand was restricted. c. 1800–1888: the ice margin at isvand receded c. 2.5 km, presumably with large variations. the rate of recession was reduced towards the end of the period. 1888 – c. 1920: the decreasing recession rate during this period shifted to a marked advance of the glacier margin at some time well before 1920, and ended with the culmination of the advance c. 1920. this may have been due to low average annual temperatures in west greenland. hence at nuuk, a period with average annual temperatures below –2°c occurred during the period from 1873 to 1920 (cappelen 2005). other factors could also have influenced the mass balance of the glacier, but it is possible that the glacier did not respond until around 20 years after the onset of cooling. pronounced moraines referred to the 1920 stade can be followed for 16 km from isvand to the glacier front, which at that time was located in kangersuneq fjord, c. 12 km behind the little ice age maximum position (fig. 12). the location of the glacier margin in isvand between the positions in 1888 and 1936 also implies an approximate culmination of this event c. 1920. 1920–1985: this period shows an increasing rate of recession beginning in 1920 and culminating c. 1955–1968 with a rate of c. 70 m per year. the recession rate fell to 12 m per year between 1968 and 1985. this pattern may reflect the temperature history in the region (cappelen 2005). the temperature record at nuuk shows that there was a broad maximum in mean annual temperatures around 1930, followed by slowly decreasing temperatures until c. 1960. the time period from 1960 to 1990 was characterised by fluctuating low temperatures. 1985–2010: the distance between the outer part of the trimline zone and the ice margin is rather stable at c. 5.5 km from the little ice age maximum extent (fig. 9b). satellite images indicate that the shrunken isvand lake in the first decade of the 2000s have drained through kangiata nunaata sermia. this happened on two occasions, around 2004 and 2009. thus after at least c. 250 years, the drainage of lake isvand to ameralla has changed back to the medieval conditions of the area with a very modest water discharge through austmannadalen (weidick & citterio 2011). the kangiata nunaata sermia glacier system conditions around the head of kangersuneq fjord after the time of the norse settlers cannot be described better than by roussell (1941, p. 16–18), who after a description of ameralik and austmannadalen writes about kangersuneq: this is the great fjord, now called kangersuneq, which from the innermost part of the settlement cuts its way southeast towards the inland ice and embraces three glaciers, one of which is dead whereas the other two are very active, even if their calves are only small in comparison with those of north greenland. all the same, their production is large enough to prevent the fjord from being navigable even to small craft. it is said that the chances of getting through are best in the month of august; one year at this time we succeeded with great difficulty in forcing a boat half way up the fjord, where we had to go ashore and continue on foot. another year we arrived a little earlier, but the fairway was closed. considering that along the east shore there are several farms, and not the smallest either, it must be assumed that conditions in the middle ages were different. on the other hand one cannot be too careful when summing up natural conditions in greenland. in the same fjord it is clearly to be seen that the principal glacier once 2424 extended much farther out (fig. 8); the lower part of the rock sides is as it were divided into two parts along a line which slants evenly down from the present upper edge of the ice to the water about 10 km down the fjord. above this line the fells have the usual brownish-green vegetation, but below it they are bare and barren, as if the ice had left them only recently. now, there are ruins inland behind the place which would then have been the front margin of the glacier, which seems somewhat incredible; can it be, then, that in the time between 1360 and now the glacier advanced and then retreated? a photograph taken by rink (national museum archives) in the 1850s shows that the conditions were exactly as they are today. and one is tempted to say that if eighty years cannot spread the vegetation, the situation may just as well be eight hundred years old. the fjords then as now were the main lines of communication for the settlements, and no one left the waterside of his own free will. nevertheless, in both the vatnahverfi of the east settlement and to some extent in the west settlement there are farms lying several hours’ journey from the landing place (fig. 9), and accordingly this must be taken as evidence of the extraordinary intensity of the settlement. in the west settlement there is the possibility that these inland farms were those of the last remnants of the tribe, who in these remote regions tried to find peace from the eskimos. this, however, has been disproved with complete certainty by the excavations in the austmannadal. the fairly rich finds of artefacts were of exactly the same types as those unearthed in the main farms by the fjord, and the bones in the refuse heaps show that even inland the large marine mammals were an important economic factor, which means that the farmers in the valley cannot have concealed themselves, but must often have gone whaling and sealing out through the fjord to the open sea. the photograph mentioned as fig. 8 in this quotation shows ruin group 16 and the trimline zone on akullersuaq semi-nunatak, seen from the north. the photograph by rink mentioned in the text is also shown here (fig. 13a). roussell’s fig. 9 shows an inland norse farm in the innermost part of austmannadalen. the eastern settlement refers to the norse settlements in south greenland and the western settlement to the norse settlement in the godthåbsfjord region. little is known about the innermost parts around kangiata nunaata sermia from the 18th century. on thorhallesen’s map from 1776 (fig. 10; thorhallesen 1914) two large outlets are seen coming from the inland ice and calving in kangersuneq. the northern of these outlets must be narsap sermia and the southern is the kangiata nunaata sermia glacier system, consisting of kangiata nunaata sermia and akullersuup sermia. k.l. giesecke described the area around the head of kangersuneq, which he visited in august 1808, mentioning the place names nunatarsuaq and illorsuit. he may have camped between illorsuit (ruin group 13a; fig. 7) and ruin group 15 (at umiivik). giesecke (1910, p. 150) described the ruins as stone heaps, traces of stone walls and single buildings, covered with soil, grass and bushes. the following description of the norse ruins at saqqarsuaq is more detailed (ruin group 16, figs 3, 7). on older maps, ruin group 16 is located on the northwestern slope of the valley between qamanaarsuup sermia and the inner part of kangersuneq, c. 5 km from the icefjord (figs 3, 7). roussell (1941, p. 78, fig. 54) and kort & matrikelstyrelsen (1993), however, placed ruin group 16 near the southern tip of the semi-nunatak nunatarsuaq, near the coast of kangersuneq. according to his diary for 7–12 august 1808, giesecke (1910, pp. 147– 151) passed narsap sermia and narsaq on his way to the head of kangersuneq and then camped close to narsaq due to bad weather. the following attempt to reach the front of kangiata nunaata sermia glacier system took place on 12 august, when giesecke and his companions arrived at the illorsuit site (ruin group 13a, figs 3, 7), described as being located at the end of kangersuneq. they reached ruin group 16 (saqqarsuaq) in the valley between qamanaarsuup sermia and the frontal parts of kangiata nunaata sermia. giesecke wrote (pp. 150–151) the following about the valley, the glaciers and the ruin group: this valley runs to the opposite side back down to the sea, and has as the former in its centre a large raging mountain river. part of it runs straight to the sea, and part of it runs via another curvature around a small mountain towards the glacier to the east. this glacier is larger, steeper and more dangerous than the one to the north-east, and there are no nunataks, one sees only the immense masses of towering ice. here, very close to the ice, i found a well-preserved norse ruin that protruded above the small shrubs. it formed a square and measured close to 50 × 50 feet, and in some places the walls reached c. 1 yard [1 m] above the moss, grass and shrubs. the inner walls were difficult to discern, but the shrubs indicated three main rooms. several other buildings in this rocky valley were probably buried below the glacier long ago. i ascended both mountains located between the described valleys in the area. these mountains are extremely fissured; the fissures are sometimes 1 fathom [2 m] or more wide and very long. the surface of the mountains is covered by enormous rock fragments, which indicates a fairly recent 25 earthquake, because the rock fragments, which here and there fit together are still very sharp. on the other side of the valley, towards the glacier, a large lake is found, which drains over a bedrock threshold to the glacier and on to the sea. in the valley near the lake some reindeer grazed quietly until i was around fifty paces away... giesecke described the view from the mountain (c. 1000 m above sea line): from the higher of the aforementioned mountains i could see pisissarfik, a part of ameraglikfiord with its small islands, as well as both glaciers in this region. the glaciers actually belong to the great ice cover that is found on all land from north-west to south-east. the sight of this ice cover is incredibly beautiful, and well worth the difficult ascent. (authors’ translation.) ‘pisigsarfik’ was the fjord that is now called kapisillit kangerluat; the name pisissarfik is now used for a 1220 m high mountain peak on the northern shore of this fjord. following the descriptions above and a later description by roussell, the location of ruin group 16 near kangersuneq can be determined. the presence of a large ice-dammed lake filling the valley in front of qamanaarsuup sermia must imply an advanced position of the kangiata nunaata sermia system. the general description also indicates that the front was close to ruin group 16, and that the glacier dammed the meltwater from qamanaarsuup sermia to form an ice-dammed lake (called ‘saqqarsuaq lake’ in the following). knuth (1944, p. 100) doubted that giesecke had visited ruin group 16, which he suggested should be located farther to the east in accordance with d. bruun’s map (fig. 7). however, roussell’s description of ruin group 16, and roussell’s localisation of ruin group 16 to a position close to the eastern margin of kangiata nunaata sermia based on the description of giesecke, fits with roussell’s own measurements of the ruin group. kleinschmidt’s map from 1860 (fig. 11) shows the front of kangiata nunaata sermia in a more retracted position than described by giesecke in 1808, when kangiafig. 13. a: the front of kangiata nunaata sermia, seen from the north. the photograph was taken in the 1850s by h. rink; it is housed in the archives of the national museum in copenhagen. b: same as above, taken by john møller in 1903. reproduced from bruun (1917, fig. 11, p. 76). the location of the photographer was not exactly the same as for a. h. rink (1819–1893) was a royal inspector of south greenland and later a director of the royal greenland trade. in addition to his administrative duties, he is known for his pioneering studies of the geology, glaciology and anthropology of greenland. j. möller (1867–1935) was born in greenland and educated in denmark as a printer, photographer and interpreter. b a fig. 13 akullersuaq nunataarsuk 1850s 1903 2626 ta nunaata sermia blocked the valley to qamanaarsuup sermia leading to the formation of an ice-dammed lake. on kleinschmidt’s map the glacier front is placed south of the akullersuaq semi-nunatak, 10–12 km from the little ice age maximum position as shown by the trimline zone. the same position is shown on j.a.d. jensen’s map from 1885. the earliest photograph of a glacier in greenland may be that taken by rink in the 1850s. rink’s photograph shows the front of kangiata nunaata sermia, seen from the west side of kangersuneq at the isthmus between kapisillit kangerluat and kangersuneq, north-east of the kapisillit settlement (fig 6b). rink’s photograph, which is also mentioned by roussell (1941, p. 17), is reproduced here (fig. 13a). a later photograph taken from approximately the same position by john møller in 1903 and published by bruun (1917, fig. 10, p. 76) is shown here as fig. 13b. the trimline zone on the mountain walls of the semi-nunataks nunataarsuk and akullersuaq can be seen and the position of the glacier front on both photographs can be estimated to 10–12 km south of the little ice age maximum extension of the glacier. following a visit in 1909, nordenskiöld (1914, p. 638–639) wrote that kangiata nunaata sermia had been retreating, in contrast to narsap sermia which was advancing during the first years of the 1900s. concerning ruin group 16 at saqqarsuaq, bruun (1917, p. 78) wrote that he was told by greenlanders during his visit in 1903 that there was a rather small ruin, but earlier there could have been more ruins, which may have been flushed away by a river or destroyed by the ice. both bruun’s and nordenskiöld’s evidence for recession of kangiata nunaata sermia could refer to the middle or the last decades of the 1800s, and the advance leading to the 1920 stade may have started in the last decades of the 1800s or the beginning of the 1900s. the older little ice age maximum advance seems therefore to have occurred in the 1700s and giesecke’s description of a lake in the valley between qamanaarsup sermia and the kangiata nunaata sermia glacier system implies a situation with the front of the kangiata nunaata sermia glacier system close to the little ice age maximum extent. bruun’s record of other ruins at the site of ruin group 16 being removed by the nearby glacier or flushed away by its marginal drainage is easy to understand if it occurred during the maximum extent of the kangiata nunaata sermia glacier system. according to the trimline zones, the glacier in this area at its little ice age maximum reached an elevation of 123 m a.s.l. (roussell 1941, p. 79). the present ruins are located at 208 m a.s.l., according to roussell’s description of the ruin site (roussell 1941, fig. 54, p. 78), which is indeed close to the trimline zone (fig. 14). the apparently fairly stationary ruins kangersuneq 1 km 100 m 200 m fig. 14fig. 14. ruin group 16 and the surrounding terrain at saqqarsuaq, kangersuneq icefjord. redrawn from roussell (1941, fig. 54). the red part of the terrain indicates the estimated extent of the ice cover of kangiata nunaata sermia at its maximum during the little ice age. the slope of the trimline zone maximum is c. 20 m/km, decreasing towards the north. according to roussell (1941), the altitude of the trimline zone is 123 m a.s.l., and the ruins are at 208 m a.s.l., although the latter figure is not in accord with the contours on the map. note that this is a reproduction of the original map. facing page: fig. 15. a: front of kangiata nunaata sermia in 1921 seen from the west. photograph by aa. nissen in geus’ glaciological archive (filed under glacier 1ch23003); aage nissen (1889–?) was a teacher at and later a leader of the teacher’s college in nuuk in the period from 1920 to 1927. b: oblique aerial photograph of the interior part of kangersuneq looking from the south-west over kangersuneq to the north-east, where the front of qamanaarsuup sermia can be seen, surrounded by a broad trimline zone. in the foreground a double set of marginal moraines can be seen. the outer set (labelled lia max) marks the extension of the kangiata nunaata sermia glacier system in the 1700s. the younger set (labelled 1920 stade) marks the maximum of a large readvance, presumably initiated during the beginning of the 1900s and culminating shortly before 1920. geodetic institute, route 506b-n, no. 6270, 21 august 1948. 27 b a 1920 stade lia max kangersuneq qamanaarsuup sermia fig. 15 1948 1921 2828 glacier front at the end of 1800s and beginning of the 1900s astonished roussell (see quotation on pages 2324). he estimated the front of kangiata nunaata sermia to be c. 10 km from its little ice age maximum position, indicating a century of ‘stability’ of the position of the glacier front. however, the large fluctuation during the 1920 stade must have taken place during this period. with respect to ruin group 15 (umiivik), bruun (1917, p. 78) referred to the statement of the greenlanders, viz. that it is a locality with many, partly overgrown ruins, which is in accordance with information from a visit to the place by o. bendixen in 1916 (mentioned in bruun 1917, p. 72). as mentioned above, j.a.d. jensen’s map was widely used up to the 1930s. a photograph from 1921 shows the front of kangiata nunaata sermia seen from the north-west (fig. 15a). the glacier front is located close to the moraines that formed during the 1920 stade. during the first decades of the 1900s, ice-margin features formed in the central parts of the trimline zone (fig. 15b). these ice-margin features are widespread and dark in several places, which imply dead ice or ice-cored moraines. traces of this advance (the 1920 stade) are also seen at isvand. 50°w 50°w 51°w 51°w 65 °n 65 °n 64 °3 0' n 64 °3 0' n 64 °n 64 °n kangaasarsuup sermia kangersuneq austmannadalen nunataarsuk qamanaarsuup sermia akullersuaq saqqarsuaq (16) umiivik (15) nunatarsuaq tummeralik (37) lia max lia max isvand austmannadalen 1920 1936 2009 198 5 akullersuup sermia 1948 kangiata nunaata sermia qamanaarsuup sermia lia max akullersuaq nunataarsuk saqqarsuaq (16) umiivik (15) nunatarsuaq tummeralik (37) lia max 1808? 2 km 1920 lia max 1968 64°15´n64°15´n langvand 64°15´n 49°30´w50°w 49°30´w fig. 16. recession of kangiata nunaata sermia since the little ice age maximum (lia max) in the 1700s. the oldest ice margin known after the lia max is the presumed position of the glacier front in 1808, when the glacier prevented water from draining from the qamanaarsuup sermia glacier to the kangersuneq fjord. during later positions from c. 1850 to the 1930s, the front was located somewhere between the southern tip of the akullersuaq seminunatak in the north-east and the area eastnorth-east of tummeralik (ruin group 37) in the south-west. umiivik: ruin group 15, saqqarsuaq: ruin group 16. the map is based on aerial photographs from 1985 when the ice-dammed lake isvand was at its maximum extent. 29 a more general recession began in the 1920s, but on aerial photographs from 1948 kangiata nunaata sermia and akullersuup sermia are still confluent, forming a ‘kangiata nunaata sermia glacier system’ where the glacier tongues have a common front in kangersuneq. not until the end of the 1940s was the recession so pronounced that the glaciers were split up into two separate outlets. according to a satellite image from 22 june 2009 (figs 12, 16), continued recession has brought the front of kangiata nunaata sermia to a position 21–22 km behind the position of the little ice age maximum in the 1700s. during the period from 1985 to 2010, the front of akullersuup sermia has only retreated 1–2 km. evaluation the sources from the 18th century all describe an advancing glacier. the front of calving tidal glaciers may show both annual fluctuations of up to a kilometre, with advance during the winter and retreat during the summer, and larger decadal fluctuations. the historical record can provide information about the latter, but usually only give general trends, and only rarely contribute with accurate data. the little ice age maximum extent was presumably reached for the first time in the beginning of the 1700s, but as some sources also speak about advance later in that century, it is possible that the ice margin was close to the little ice age maximum throughout the century, although with minor fluctuations: c. 1775–1810: the frontal advance of the kangiata nunaata sermia glacier system may have culminated near a headland in kangersuneq, located just south of the ruin group at umiivik (qassertup nuua). thorhallesen’s map from 1776 gives little information but the glacier at the head of what is called ‘ujaraksoak fjorden’ on the map must be the kangiata nunaata sermia glacier system. this map shows the length of kangersuneq fjord from the mouth at karra (fig. 3) to the head of the fjord at the little ice age maximum of the kangiata nunaata sermia glacier system to be about 6 miles; since 1 danish mile is 7.5 km (norsk forlishistorisk forening 2012; marcussen 2011), this equates to c. 45 km, a distance that is compatible with modern maps (e.g. see fig. 3). the overall advance ended in the 18th century, and was followed by a slow recession that may have continued to the beginning of the 19th century, when the kangiata nunaata sermia glacier system still blocked the valley to qamanaarsuup sermia. the ‘saqqarsuaq ice-dammed lake’ existed in the beginning of the 1800s, implying that c. 1808 the glacier front was still close to the maximum extent of the little ice age, which can be mapped from the well-defined trimline zone (fig. 14). 1810–1860: during the following decades, fast recession took place, so that the front at the latest around 1860 had retreated c. 10–13 km from the little ice age maximum extent. during this c. 50 year period of recession, the average recession rate was 200–260 m per year. the information from giesecke about an ice-dammed lake (the ‘saqqarsuaq lake’) between kangiata nunaata sermia and qamanaarsuup sermia in 1808 and the short distance (c. 5 km) from the little ice age maximum of the frontal parts of kangiata nunaata sermia to the ‘saqqarsuaq lake’ must imply a glacier dam by a large ice body (fig.17a). roussell (1941, p. 79) measured the maximum height of the trimline zone at ruin group 16 (saqqarsuaq) to have been 123 m a.s.l., i.e. close to the lake. the exact lake level is unknown, but it was presumably c. 48 m a.s.l. (fig. 17a). the damming ice body can scarcely have had a much smaller thickness than at the little ice age maximum, and we therefore suggest that the glacier front in 1808 still had a position close to the maximum of the glacier system. as mentioned in the introduction, depths over 600 m are found in the inner part of nuup kangerlua. off kapisillit, the bottom of kangersuneq is flat at c. 340 m, and no indication of a submarine threshold or moraine is observed that may have been formed during the kapisigdlit stade. in contrast, ice-margin deposits from the little ice age maximum form a threshold at 150 m below sea level (b.s.l.). the 1920 stade is only marked by a small rise of the sea bottom to a depth of c. 300 m. it is possible that both thresholds were anchoring points for the front for longer periods, but it is not known if the thresholds consist of bedrock or ice-margin deposits (fig. 8). the increase of ice in kangersuneq before 1885 (j.a.d. jensen 1889, p. 88) could have been connected to a large breakup during the thinning of the outermost part of the kangiata nunaata sermia glacier system. as large calf-ice production may be a sign of a retreating glacier rather than an advancing glacier, the increased calfice production could indicate a period with fast recession in the first half of 1800s. 1860–1903: information from 1850 and the following decades are based on the map of kleinschmidt, and 3030 qamanaarsuup sermia a b b a b lia max 1948 1985 1968/2009 1 km 64°28´n 64°28´n 49°40´w 49°40´w former ice-dammed ‘saqqarsuaq lake’ fig. 17 1943? 2009 fig. 17. a: qamanaarsuup sermia, seen from the west. in the foreground near the front of the glacier a proglacial lake terrace (a) is seen. it is presumably related to the icedammed ‘saqqarsuaq lake’, that occupied the valley between qamanaarsuup sermia and kangersuneq until the beginning of the 1800s, as observed by k.l. giesecke in 1808 (giesecke 1910). b–b marks a moraine ridge that may be referred to the kapisigdlit stade (tentatively dated to 8.2 ka bp). aerial photograph, geodetic institute, route 505 e-ø, no. 7418 (probably from 1943). b: qamanaarsuup sermia with frontal positions at the little ice age maximum, in 1948, 1968 and 1985 according to taurisano (2004, fig. 3.5), plotted on an aster image from 2009. rink’s and møller’s photographs from the 1850s and 1903. they indicate that the glacier front, at least at these specific times, was located at a position 10–12 km from the little ice age maximum, as also suggested by aa. roussell. 1903–1920: only little information is available from the first two decades of the 1900s. although it could seem that the front was essentially stable from c. 1850 to c. 1920, however, significant changes in the frontal position must have occurred during this period. a pho31 tograph by aa. nissen from 1921 (fig. 15a) shows the front of the kangiata nunaata sermia glacier system. the glacier front was close to, but behind the moraines of the 1920 stade, and the 1921 glacier front was closely surrounded by fresh moraines deposited just prior to 1921. the moraines are the result of a marked advance, presumably initiated one or two decades before. the icemargin features can be seen in the trimline zone on the west side of kangiata nunaata sermia, extending from an altitude of around 400–450 m a.s.l. in the isvand area (c. 15 km from the front) and down to the calving front in kangersuneq fjord. a large initial recession must have taken place before c. 1900, but how far up the kangersuneq fjord is not known. the following advance probably occurred during the first two decades of the 1900s. a clear colour difference is seen between the trimline zone below and above the line of the 1920 stade on photographs and satellite images (fig. 9b). the upper and older part of the trimline zone, which is from the 1700s, is characterised by widespread lichen and vegetation cover, as well as some soil development. the lower and younger part of the trimline zone was not left finally by the glacier until after c. 1920, and this part is therefore more barren than the older part. however, both parts can be characterised as ‘barren ground’ that is clearly different from the surrounding areas with older vegetation. 1920–1985: a period of fast recession of the front took place after c. 1920. around 1950–1960, this led to a separation of the kangiata nunaata sermia glacier system into the kangiata nunaata sermia proper and akullersuup sermia, which is now a separate outlet from the ice sheet. the later development of the frontal position of kangiata nunaata sermia is shown by the following estimated positions of the front behind the little ice age maximum: 1920: 11.5 km, 1921: 11.8 km, 1936: 14.5 km, 1948: 18 km and 1968: 20.5 km (fig. 16). an advance took place in the following years and in 1985 the front had advanced to a position c. 20 km behind the little ice age maximum, as also pointed out by timm (2010). 1985–2010: by 2009, the front of kangiata nunaata sermia reached a position 22 km behind the little ice age maximum. according to a 2010 aster image, however, it had advanced to a position c. 20.5 km from the little ice age maximum. these secondary fluctuations of the frontal position during the last decades presumably reflect minor annual fluctuations of the glacier, the frontal position of which is determined by anchor points such as submarine thresholds or other features of the fjord topography (e.g. mercer 1961). regarding the recession of akullersuup sermia, we note that the height of the trimline zone above this glacier also indicates a marked thinning, as recorded for kangiata nunaata sermia. after the split of the glacier system, the front of akullersuup sermia has receded only c. 2 km between 1985 and 2010. qamanaarsuup sermia this glacier has already been mentioned above in connection with the formation of the ice-dammed ‘saqqarsuaq lake’ in front of the glacier (fig. 17a). its more recent development is described below. the glacier and its surroundings are located in the most remote part of the kangersuneq region. although its name refers to reindeer hunting (qamavoq = wait for it), it must have been difficult to access during the little ice age maximum in the 1700s when the valley in front of it was closed by the ‘saqqarsuaq lake’ and the kangiata nunaata sermia glacier system. hence the first information about this glacier and the valley in front of it was published following the mapping by j.a.d. jensen in 1885 (jensen 1889). the glacier is now surrounded by a trimline zone, the outermost parts of which appear to be related to the proglacial, ice-dammed lake that may have existed in the 1700s. morphological evidence of a former ice-dammed lake was also reported by andsbjerg (1985). this lake may have drained in the first half of the 1800s, but a more complete understanding of the history of this lake must be left to future investigations of the lacustrine deposits in the valley. the oldest aerial photograph that covers qamanaarsuup sermia dates from 27 august 1936 (geodetic institute route 61a, no. 25283, not shown). unfortunately, this oblique photograph was taken from a distance and the glacier lobe can only be faintly seen surrounded by a trimline zone. taurisano (2004) puts the last advance to the little ice age maximum at 1880–1890 on the basis of the lichenometrical dates of beschel (1961). however, beschel’s data do not preclude that the maximum position was reached earlier, in the 1700s. beschel’s data on growth rates of lichens are from central west greenland, far from the kangersuneq region, and are mainly based on fluctuations of local, small glaciers. subsequent frontal positions 3232 of qamanaarsuup sermia between 1948 and 1993 were mapped by taurisano (fig 17b; taurisano 2004, fig. 3.5). from the little ice age maximum to 1948 the glacier receded 1.5 km, corresponding to an average value of 24 m per year, if recession began in 1885 as suggested by taurisano. from 1948 to 1968, the glacier receded 0.5 km, corresponding to 25 m per year. between 1968 and 1985, recession was very limited. from 1985 to 1993, the glacier front receded c. 100 m, corresponding to 13 m per year. finally, on a landsat image from 2009 the front was at approximately the same position as in 1993. on the basis of aerial photographs from 1968 and 1980, detailed maps (1:10 000, contour interval 10 m) were made by knudsen (1983). from a comparison of the two maps, knudsen concluded that the glacier had receded 25–50 m and he recorded a thinning near the snout of the glacier to less than 10 m. the contour lines at higher elevations were nearly identical in the two years. evaluation little is known about the fluctuations of this glacier before the 1900s. its frontal deposits at the former icedammed lake may indicate that the maximum extent of the glacier and the formation of the ice-margin deposits took place at the same time as the possible formation of the lake, so that the history of this glacier may be similar to that of kangiata nunaata sermia. however, the later recession of this land-based glacier so far only amounts to around 2 km. during the more detailed measurements of the changes of this glacier between 1968 and 1985, the recession was only 25–50 m, i.e. 1–3 m per year. the presumed ice-dammed lake (‘saqqarsuaq lake’) must have been emptied during the recession between 1810 and 1850. the terrace in fig. 17a marked ‘a’ may be a remnant of the alluvial plain that formed during the little ice age maximum; the surface of the terrace would thus mark the level of the former ‘saqqarsuaq lake’ at its maximum extent. the elevation of the terrace was measured by gps in the summer 2011 to be 48 m a.s.l. (dirk van as, personal communication, 2011), which must indicate the maximum elevation of the lake level. because the elevation of the former lake bottom is 0–30 m a.s.l., the lake must have been shallow even though it had an extent of c. 10 km2. observations of the frontal positions of qamanaarsuup sermia in 1948, 1968 and 1993 (taurisano 2004) show positions of 1.5, 2.0 and 2.1 km behind the little ice age maximum, respectively, and indicate that the recession rate began to decrease in the 1960s, possibly induced by a minor cooling of the climate. according to a landsat image, in 2009 the frontal position was close to that of 1993, as mapped by taurisano (fig. 17b). according to taurisano (2004, pp. 36–39), the lower and upper parts of qamanaarsuup sermia showed marked differences in surface elevation changes from 1968 to 1985. the upper part thickened by 1 m per year whereas the lower part thinned by 0.3 m and is gradually being transformed into dead ice (see frontispiece). narsap sermia the surroundings of narsap sermia have been described several times due to the occurrence of norse ruins (groups 11 and 12; figs 3, 7) near the front of narsap sermia (figs 3, 7, 18). early records of these ruins are from bruun (1917, pp. 75–77), who included information from o. bendixen from a journey in 1916 and from greenlanders who noticed these ruins close to the front of the glacier. narsap sermia was plotted on the early map by thorhallesen (fig. 10). although his map is sketchy and does not depict ruin sites in this area, the glacier front is, as today, shown at some distance from kangersuneq. this is in contrast to the more northern fjord, on his map called ‘naviangoæt’, where the norse ruin groups 7 and 8 are located (figs 3, 7). the proximity of the front of narsap sermia to kangersuneq and to narsaq was also recorded by giesecke (1910, p. 147–148). on 7 august 1808, he passed its front on the way from ujarassuit paavat to narsaq (ruin group 12, figs 3, 7), and continued to a place called ‘narkseeitsiaq’, where he camped. the camp site must have been relatively close to narsaq since he described calving noise from the front of narsap sermia. j.a.d. jensen’s map, based on observations in 1884 and 1885, shows a position of the front closer to kangersuneq. in his description of ruin groups 11 and 12, jensen does not mention the nearby front of narsap sermia (jensen 1889, p. 110). from a visit in 1909, nordenskiöld (1914) reported that narsap sermia had advanced during the first years of the 20th century; the advancing glacier had buried several hunting camp sites. d. bruun (1917, p. 75) noted, when describing ruin group 11 (saqqannguaq), that ruins are found at the 33 shores of rivulets. this is based on observations by o. bendixen from his voyage in 1916. for ruin group 12 (narsaq), d. bruun (1917, p. 77) reported that at narsaq, between a steep headland and the end of the glacier, there is a small ruin on the slope immediately above the fjord. some of the ruins had slipped into the water. the descriptions of roussell, after visiting the area in the 1930s, give no new information about narsap sermia. aerial photographs and satellite images from 1936 onwards show that the front of narsap sermia has been located at about the same place. it appears that the front of this glacier may have been relatively stable for around 250 years, although this conclusion is based on scattered information. general recession of the glacier front seems to have started in the first decade of this century. evaluation it appears from the map of thorhallesen (1914; fig. 10) that the front of narsap sermia was situated in a small fjord branch, a short distance from kangersuneq. the same is seen on aerial photographs from 1936, 1948 and 1985, as well as on satellite images from 2009 and 2011. the glacier is only surrounded by a narrow trimline zone, which supports the proposal that it has not fluctuated much during the past centuries. based on studies of aerial photographs and satellite images, timm (2010) considered the glacier stationary or advancing between 1936 and 1985, but it was suggested that the central parts of the glacier front receded c. 700 m between 2002 and 2009. this is in accordance with a narrow trimline zone developed at the lower flanks of the glacier. it appears to be the first signs of the following frontal recession of around 1 km that occurred in the period 2009–2011 (landsat image, 14 july, 2011), as measured along the centre line of the glacier (fig. 18). fig. 18. narsap sermia. the frontal position of the glacier is taken from a landsat image from 2009 and drawn on an aerial photograph from 16 august 1968, geodetic institute, route 281t, no. 70. in 1968, the norse ruins at narsaq and saqqannguaq were located around 1 and 2.5 km from the glacier front, which is the local little ice age maximum of the glacier. marked recession of the centre line of the glacier (c. 1 km) occurred between 2009 and 2011. 64°40´n 1968 2011 2009 1 km saqqannguaq narsaq narsap sermia ivisaartoq 64°40´n 64°40´n 50°05´w 50°05´w fig. 18 1968 3434 fig. 19. a: kangilinnguata sermia on 25 august 1930, seen from a height of 630 m a.s.l. (gripp 1975, plate 2, fig. 2). the arrow indicates the location of norse ruin group 8 (figs 3, 7), which was mapped, described and photographed with the glacier in the background by roussell in 1934 (roussell 1941, pp. 76–78). b: kangilinnguata sermia seen from the north in 1948 (geodetic institute, route 505-es, no. 1948). the lake south of the glacier front (top left) receives large amounts of meltwater from the inland ice margin south of kangilinnguata sermia. this meltwater drains to the ujarassuit paavat fjord, to the right (see fig. 3) via the outermost part of the front of kangilinnguata sermia. a naajat kinginnerat kangilinnguata sermia narsap sermia ivisaartoq ujarassuit nunaat b kangilinngua naajannguit? 1930 1948 35 kangilinnguata sermia in contrast to the other outlets around kangersuneq described above, the interior of ujarassuit paavat was described by hans egede, who lived in greenland from 1721 to 1736. his descriptions of the ujarassuit paavat fjord region concentrated on a large group of ruins 8 km from the glacier (ujarassuit; ruin group 7, figs 3, 7), whereas only few persons have described the more modest ruin group at puilasoq (group 8, figs 3, 7), situated only c. 2 km from the glacier front. this is because a c. 7 km long tidal flat separates the navigable outer part of ujarassuit paavat from the inner part of the fjord with the puilasoq ruin group 8 and the kangilinnguata sermia glacier. e. thorhallesen seems to have visited the inner parts of ujarassuit paavat. on his map from 1776 (fig. 10), he calls this fjord ‘naviangoæt’ and the ujarassuit site is designated ‘europ. rudera’. the un-named ruin group at puilasoq is marked by a single building. the ice-sheet margin north-east of puilasoq is marked ‘jökull’. in his description of the locality, thorhallesen (1914, pp. 46– 47) wrote: at high tide this bay is twice as long, and from a high mountain i saw three or four more rudera [ruins] in the area. the ruins can be seen from a great distance, because the grass, where the field has been fertilised, dies and becomes red in late summer, since it starts growing earlier in the spring. the glacier ice, which covers the uplands from north to south, as far as can be seen, and which covers the highest mountains, falls towards the south-east from naviangoæt down to the sea, which is perhaps the head of ujaraksoak fjord, unless it continues farther below the ice. sometimes large pieces of ice fall down and float around until a strong eastern wind drives them out of the fjord. (authors’ translation.) although the frontal position of kangilinnguata sermia is not well defined, it seems that the front during the last half of the 18th century was advancing and that the glacier calved into the fjord, at least at high tide. hence it is possible that the tidal flat in the inner part of the fjord formed after thorhallesen’s visit. k.l. giesecke also visited ujarassuit paavat. in his diary, he mentioned on 5–6 august 1808 (giesecke 1910, pp. 145–147) the earlier visits by h. egede and e. thorhallesen, but he only briefly mentions the glacier. kleinschmidt’s map from 1860 (fig. 11) depicts ujarassuit paavat, the bay at ujarassuit and the head of an outlet from the inland ice that almost reaches the inner part of ujarassuit paavat. the latter outlet must be kangilinnguata sermia. the sketchy contours of the map do not allow us to determine the exact position of the glacier front, but it appears that the glacier did not reach the sea. after a later visit in 1885, by j.a.d. jensen (jensen 1889, p. 109), ruin group 8 and its surrounding area were described thus: should be located on the north side of ujaragssuit’s north-eastern branch, into which the inland ice now protrudes, transporting large amounts of clay that are deposited in the fjord, so that it becomes dry at low tide. living greenlanders remember that until around 1840 they could sail up this fjord branch by umiaq to a bird cliff on the south side, but the ice has advanced far beyond this place. (authors’ translation.) based on the statements by thorhallesen from the middle of the 1700s and those by j.a.d. jensen from the middle of the 1800s, we suggest that kangilinnguata sermia was advancing at this time and that sedimentation in ujarassuit paavat was increasing. although visited subsequently by d. bruun in 1903 (bruun 1917) and by o. bendixen in 1916 (both mentioned in bruun 1917, p. 72), no further information was provided about the glacier. aa. roussell, after work in 1934, provided detailed descriptions of ruin group 7 (anavik, now ujarassuit; roussell 1941, pp. 32– 34) and group 8 (puilasoq, pp. 75–77, 167–171). roussell stated that the distance from ruin group 8 to the front of kangilinnguata sermia was around 2 km (roussell 1941, p. 76). he published a photograph showing the glacier at approximately the same position as at present (roussell 1941, fig. 52, p. 76). the glacier was also described by gripp, who investigated the glacier in 1930 and described marine faunas found in concretions in the moraine and in the alluvial plain in front of the glacier (gripp 1932). gripp also described the geomorphology of deposits in front of the glacier and within the glacier, where push and shear moraines occurred (gripp 1975). gripp (1932, 1975) published photographs of kangilinnguata sermia (fig. 19a). it appears that the glacier still had no pronounced trimline zone around 1930, but the glacier front was heavily crevassed and partly dominated by shear moraines and about 10 m high push moraines close to the active glacier ice. gripp’s photographs can be compared with aerial photographs of the glacier from 1948, 1968 and 1985 as well as a satellite image from 2007 (figs 19, 20). apart from a slight increase in the size of the superficial moraine area in 1948, only small changes are seen in the frontal appearance, in addition to a little thinning. 3636 evaluation thorhallesen’s description from the 1700s indicates that kangilinnguata sermia was calving in the fjord now named ujarassuit paavat, but called ‘naviangoæt’ by thorhallesen. it is possible that the interior part of ujarassuit paavat was not filled with sediments to such a high degree as at present; if so, kangilinnguata sermia may have been a calving glacier even if the front was located at the same place as today. however, it is also possible that calf ice drifted into ujarassuit paavat from the kangersuneq icefjord. from the descriptions from the 1700s and most of the 1800s, it is difficult to determine the position of the glacier front. the records state that the glacier was advancing and possibly at the same time filled the interior parts of the fjord with sediments. since only a narrow trimline has developed around the glacier front, its position cannot have been farther west than today during the little ice age (including the narrow belt of shear moraines at the present glacier front). the formation of the proglacial tidal flat can only be explained by increased sedimentation. this is in accordance with the record by j.a.d. jensen that the head of the fjord could be reached by boat as late as around 1840, but not at the time of his visit in 1885, due to the development of a tidal flat. the formation of this tidal flat can hardly be taken as evidence of a glacier advance, but the notes in j.a.d. jensen’s report on the glacier front passing a bird cliff may provide information about the timing and extent of the advance. thorhallesen used the name ‘naviangoæt’ for the inner part of ujarassuit paavat. according to l. bobé’s footnote to thorhallesen’s text (thorhallesen 1914, p. 46), this place name can be translated as ‘naujanguit’, which is used by the greenlanders for the glaucous gull (larus hyperboreus) and the iceland gull (l. glaucoides); ‘naviangoæt’ is thus probably a name for a bird cliff with breeding gulls. the southern side of the frontal parts of kangilinnguata sermia borders on a 500–700 m high steep mountain slope, and the mountain is called naajat kinginnerat, but no bird cliff is known from the site (d. boertmann and l.m. rasmussen, personal communication 2011). however, this could be caused by the present inaccessibility to the place because of the advancing glacier. confirmation fig. 20. a: kangilinnguata sermia on 25 august 1968 (geodetic institute, route 281v no. 810). b: kangilinnguata sermia front on 4 august 2007 (aster image). comparisons between the photographs from 1930 and 1948 (figs 19a, b), the aerial photograph from 1968 and satellite images from 2007 and 2009 indicate a slight thinning. in addition, a thin trimline zone developed around the glacier front during this 77 year period. lia max (little ice age maximum). kangilinngua lia max 50°w 1 km kang ilinn gua ta s erm ia naajat kinginnerat lia max 1968, 2009 a b 64°50´n 64°50´n 50°w 49°50´w kangi linng uata ser mia 1 km fig. 20 1968 2007 37 (or the opposite) by local hunters is needed. however, both maps by kleinschmidt (1860) and jensen (1889) give the impression that the glacier front in 1885 was still close to its little ice age maximum extent. according to the information presented above, the glacier must have advanced in the first part of the 1800s, after which thinning and development of superficial moraines have occurred throughout the 1900s (fig. 20). it is deemed likely, therefore, that j.a.d. jensen’s reference to the advancing glacier passing a bird cliff around 1840 refers to the steep southern mountain slope, today situated 1–2 km behind the little ice age maximum as determined from terminal moraines and a narrow trimline zone. it is suggested that the glacier front passed the bird cliff in the first part of the 1800s during an advance that reached the little ice age maximum extent sometime at or after the middle of that century. the magnitude of the advance may have been between 1 and 2 km (the distance between the cliff and the little ice age maximum extent; fig. 19). the first accurate evidence of the position of the glacier front was provided by gripp after his visit to the glacier in 1930. from his description (gripp 1932, 1975), the glacier seems to have been surrounded by little ice age moraines and a trimline zone, and it must be concluded that the glacier front was in a stage of initial thinning. the distance from the glacier front to ruin group 8 in 1934 was estimated by roussell to be 2 km, which is close to the present distance. the puilassoq ruin site and kangilinnguata sermia are shown on a photograph from 1934 by roussell (roussell 1941, fig. 52, p. 76). aerial photographs from 1936, 1968 (fig. 20a) and 1985 give the impression that the glacier front was thinning and becoming increasingly covered with surface moraine. the front was probably receding slightly. a landsat image from 2009 indicates continuous thinning of the front with further development of the trimline zone, but development of dead ice makes it difficult to determine the exact recession of the glacier front. the same is seen on an aster image from 2007 (fig. 20b). comparisons with regional glacier fluctuations the description of the five outlets from the inland ice that are found in the kangersuneq area shows the great diversity in response to past climate changes. in order to discern a possible pattern in the geographical distribution of these responses, it was necessary to look at changes of the outlets from a larger geographical region of the western slope of the inland ice. preliminary investigations of calving glaciers in south-western greenland were made by weidick (1994a, b) and of the south-western slope of the ice sheet in general, based on aerial photographs from around 1950 and 1985 (weidick 1991a, b). it appears that thinning and recession since the little ice age are connected with piedmont-like outlets spreading over extensive lowland areas such as the qassimiut lobe (podlech 2004) and frederikshåb isblink (described here) in south-west greenland with relatively small frontal changes, or with large ice streams such as kangiata nunaata sermia and jakobshavn isbræ (sermeq kujalleq) in west greenland that show large frontal changes (fig. 21). fluctuations of the ice margin are also recorded in other parts of the ice sheet, but the amplitudes are small, and hence it is difficult to identify and date the former changes. however, it appears that a small but widespread advance was characteristic for outlets and marginal parts from around 1950 to 1985. at most localities, the margin only advanced up to a maximum of a few hundred metres during this period. the other outlets from the ice-sheet margin in the godthåbsfjord region are described below, beginning from the south. south of kangersuneq to frederikshåb isblink the outlets in this region are frederikshåb isblink, nakkaasorsuaq in allumersat fjord (bjørnesund), sermeq glacier, isortuarsuup sermia, 1cg14004 and kangaasarsuup sermia (fig. 2). un-named outlets are described using their inventory code numbers according to weidick et al. (1992). 3838 frederikshåb isblink frederikshåb isblink is a large, land-based outlet from the inland ice, with a circular piedmont lobe (definition by armstrong et al. 1973). it has a frontal diameter of c. 25 km and is separated from davis strait by a 5–10 km wide alluvial plain (fig. 22). this large outlet attracted early attention, and was described by several travellers. it was described in the 1700s by hans egede (1741, 1925), erich larssøn (1942), poul jochumsen moltzou (1935), lars dalager (1915) and egil thorhallesen (1914). however, these descriptions are often vague, and it is difficult to identify the described localities. this is also the case for descriptions by o. fabricius (1788). fabricius was a danish priest, zoologist and linguist who lived in paamiut just south of frederikshåb isblink from 1768 to 1773. he gave the first detailed description of the margin of the inland ice and discussed the origin of icebergs (fabricius 1788, pp. 69–70): the ice sheet is a most peculiar natural phenomenon, greatly exceeding the glaciers known in other countries, since it reaches from one end of the country to the other and conceals the entire interior part of the land with permanent ice, so that only some mountaintops protrude here and there, black and without ice cover. when ascending one of the highest mountains on the ice-free land near the coast but at the same time close to the ice sheet, one is faced with a frightful sight; however, one becomes eager to learn more about it. fig. 21. maps of south-western greenland with ice-sheet contours from bindschadler et al. (1989). the coloured zones show recession or growth in c. 1950 and c. 1985 (weidick 1991a, fig. 1, p. 40). the grey arrows show the major calf-ice producing outlets (jakobshavn isbræ, kangiata nunaata sermia and eqalorutsit killiit sermiat). the latter is an outlet from the eastern flank of the qassimiut lobe. ca. 1950 ca. 1985 disko jacobshavn isbræ 24 00 22 00 20 00 sønder strømfjord sarqap sermerssua isukasia kangiata nunata sermia sermilik frederikshåb isblink 0 200 km 22 00 30 00 31 00 c. 1950 c. 1985 disko jakobshavn isbræ 16 00 18 00 20 00 22 00 24 00 26 00 28 00 kangerlussuaq airport 200 km net recession & thinning net advance & growth major drains net recession and thinning net advance and growth major drains saqqap sermersua kangiata nunaata sermia frederikshåb isblink nakkaasorsuaq qassimiut lobe nuuk sermeq fig. 21 facing page: fig. 22. satellite image of frederikshåb isblink, aster pseudocolour mosaic from 1 july 2003. the altitude of the glacier north of dalager nunatakker is c. 1000 m a.s.l. the entire outlet is surrounded by a clear trimline zone indicating a general thinning of the glacier since the little ice age maximum (lia max), but the frontal recession, even near sea level, is only a few kilometres. a marked cover of superficial moraines has formed around the southern margin. the red dot shows the location of the studies conducted at qorlortoq by pitman (1973). a, b, c and d? refer to areas mapped by a. kornerup in 1878 (see fig. 23). 39 62°30´n dalager nunatakker 62°30´n 50°w 50°w 10 km dalager nunatakker qorlortoq tasersuaq lia max sioraq kangaarsuup tasersua frederikshåb isblink j.a.d. jensenj.a.d. jensen nunatakkernunatakker majorariaq c d? b a 2003 62°30´n fig. 22 4040 i believe even a superficial person will here be put into a state of profound reflection. as far as the eye can see, to the north, south or east, nothing but a sparkling plain of ice is seen. it deserves the name of an ice sea, because the ice-covered areas are situated lower than the nearest mountains in the ice-free land. this ice expands year by year more and more, growing both in height, and in extent, and hence has covered most of the land. where it meets high mountains it must stop until it grows over them, so it without hindrance can continue. it has been attempted to erect a stake on the bare land, some distance from the ice, and next year the stake was taken by the ice. so fast is its growth that greenlanders speak about localities where their parents hunted reindeer, and which are now completely ice covered. personally, i have seen trails that lead up to the inner part of the land; trails that have been made earlier but which now end in the ice, confirming the account of the greenlanders. the ice sheet advances in particular in the valleys and where these reach the sea or the heads of fjords; here it advances so much that it forms large sheets of ice on the water. (authors’ translation.) fabricius continued with a detailed description of glacier surfaces, crevasses and meltwater draining into crevasses. he also described glacier erosion, deposition of till and calving. from his detailed description it is clear that he had a thorough knowledge of the glaciers and the margin of the ice sheet in the area south of frederikshåb isblink, but it is difficult to position his described localities. it is clear from his description, however, that the ice margin was generally advancing in the latter part of the 1700s in this region. k.l. giesecke visited the area in 1809 (giesecke 1910), a. kornerup in 1878 (kornerup 1879), n.o. holst in 1880 (holst 1886) and j.c.d. bloch in 1890 (bloch 1892). from this period, the descriptions by kornerup and bloch are particularly interesting for the many details they provide about the ice and the landscape. an example of the results of the detailed investigations in 1878 is shown here (fig. 23). the map is based on surveys of selected areas of frederikshåb isblink by kornerup (kornerup 1879, pp. 132–133). it can be compared with the satellite image from 2003 (fig. 22), and with an aerial photograph of j.a.d. jensen nunatakker from 1985 (fig. 24). the rate of recession of the margin of frederikshåb isblink during this 125 year period has varied. in 1878, in the western parts of frederikshåb isblink, the glacier front was located just behind the moraines that mark the little ice age maximum (fig. 23; kornerup 1878); by 2003 the margin had receded 1–2  km. at dalager nunatakker, a thinning is seen at the lower parts of the nunataks alongside some deformation of the median moraines, shown on kornerup’s map. little change is seen on j.a.d. jensen nunatakker although some development of trimline zones is apparent. on the 1985 aerial photograph, grey snow-free areas of exposed glacier ice are seen at c. 1300 m a.s.l., south-west of the nunataks, indicating that the nunataks are situated close to the snow line. based on his visit in 1890, bloch described the western frontal, alluvial plain as a clay plain. furthermore bloch (1892, pp. 150–151) reported the presence of wide river beds testifying to high summer discharge from the ice (the expedition visited the locality at the end of may) and thick vegetation cover with large willow bushes that extended up to the moraine in front of the ice margin proper, indicating that it had not receded for a long time period. bloch (1892) reported a single main end-moraine, which was situated close to the margin of the ice. at a single locality, two moraines were recognised, spaced about 60 feet (c. 19 m) apart; the outer moraine was old, however, and already partly overgrown. locally, the ice margin was located c. 30 feet (9 m) behind the main moraine, the dimensions of which varied a great deal. at many localities, the rivers had breached the moraine. the largest moraine was 20 feet (6 m) high and 10 feet (3 m) wide; it consisted of clayey gravel with large and small rounded clasts. the margin of the inland ice was described by bloch (1892) as a smooth, descending plain; these notes are followed by a short description of the ice. both kornerup’s and bloch’s descriptions of the frontal parts of frederikshåb isblink, based on visits in 1878 and 1890, indicate a position close to the maximum for the little ice age. the oldest moraine described by bloch may have formed in the 1700s. in the following century, a map was drawn by k. gripp and s. hansen using terrestrial photogrammetry based facing page: fig. 23. detailed maps of areas in the frederikshåb isblink area (see fig. 22), surveyed by a. kornerup in 1878 (kornerup 1879). a: front of frederikshåb isblink. b: dalager nunatakker. c: j.a.d. jensen nunatakker. d: probably a small nunatak situated about 8 km east of the easternmost of dalager nunatakker, located as e on kornerup’s original map (kornerup 1879, map sheet c). 41 fig. 23 4242 on photographs taken in 1930. the unpublished map covers part of the glacier front to the davis strait on a scale of 1:10 000 and with a 10 m contour interval. aerial photographs taken between the 1930s and 1985 show a gradual thinning of the large lobe and photographs from 1985 show the trimline zone to be c. 1 km wide. at the northern part of the front of frederikshåb isblink, near the majorariaq river, the total recession from the outer moraines of the trimline zone (the little ice age maximum) to the position in 2010 can be estimated to be c. 3 km whereas the total recession of the western part of the front near sioraq is scarcely 1–2 km for the same time period. farther south at qorlortoq, at a profile of pitman (1973), the ice front shows a total recession of c. 500 m from the little ice age maximum to 2010. frederikshåb isblink is a large glacier, and it will be necessary to conduct new field investigations in order to locate and re-measure the localities described, before 1 km nakkaasorsuaq 1878 63°03´n 63°03´n 49°40´w 49°40´w allumersat li a m ax 1985 1968 fig. 25 1968 fig. 25. nakkaasorsuaq outlet in allumersat (bjørnesund). the approximate frontal positions during the little ice age maximum (lia max), in 1878 and in 1985 are shown (geodetic institute, route 281m, no. 434, 13 august 1968). the position in 2010 was the same as in 1968. 2 km 49°w 62°50´n 49°w 62°50´n fig. 24 1985 fig. 24. j.a.d. jensen nunatakker, 19 july 1985, geodetic institute, route 886n, no. 1672. when compared to fig. 23c it is seen that only minor changes have occurred since 1878, although a narrow trimline zone seen locally indicates a lowering of the inland ice surface in the region. dark areas on the inland ice surface south-west of the nunataks indicate sporadic areas of snow-free glacier ice and the proximity of the snow line at this altitude on the inland ice surface (c. 1300 m a.s.l.). the elevation of the peaks of the nunataks in this area varies from 1440 to 1680 m a.s.l. according to geodetic institute (1974). 43 a comprehensive evaluation of the changes of the entire glacier can be made. pitman (1973) carried out work on the glacial history of frederikshåb isblink near the qorlortoq locality, at 62°29.5́ n, 49°47´w, south-west of the large ice-dammed lake kangaarsuup tasersua. ages were determined using lichenometry, and the little ice age maximum was dated to 1832. between the little ice age maximum and c. 1950, the glacier has receded c. 400 m. later aerial photographs from 1964 (geodetic institute, route 272v, no. 100, 2 july) and 1985 (geodetic institute, route 886l, no. 665, 9 july) and satellite images from the first decade of the 2000s show little change in the frontal position of the glacier. this is presumably due to the strong development of surface moraine cover over the southern part of frederikshåb isblink, which has led to large areas of dead ice. nakkaasorsuaq the nakkaasorsuaq glacier in allumersat fjord at 63° 03́ n, 49°42´w is a tidal, calving glacier, but with limited calf-ice production. the front is c. 800 m wide and is surrounded by a trimline zone that extends c. 2.5 km out beyond the present front (fig. 25). early descriptions of the area date back to 1801 (mørch 1942), but give no details of the geography of the fjord, and only state that it is closed by ice (weidick 1959). giesecke visited the fjord in 1809, but he did not reach the head. mapping of the interior of allumersat started with the collection of map sketches made by local hunters (rink’s map collection in the royal library, copenhagen), but it is not possible to locate the glacier front from this source (the best example is ‘peter’s map’, mentioned by weidick (1959)). the first description and more detailed mapping of the glacier and the surrounding area took place in 1878 (jensen 1879, pp. 35–38 and his enclosed map). during the expedition, jensen visited the mountain of qaqqatsiaq c. 10 km north of the glacier on 12–14 june. after the descent, the expedition made soundings in the fjord at a distance of 0.25 mile from the glacier front and found a maximum depth of 212 fathoms. assuming danish units were used, it means that the depth was c. 400 m at c. 2 km from the glacier front (1 danish fathom = 1.8331 m, 1 danish mile = 7.5 km). the position of the glacier front in 1878 was c. 800 m behind the little ice age maximum (fig. 25). this estimate is supported by a water colour made by a. kornerup (fig. 26a) during the expedition, published by weidick (1975a) together with a later photograph from the same site, taken in 1936 by j. helk (fig. 26b). the latter photograph is the only information for the period following jensen’s expedition in 1878. it shows the front of the nakkaasorsuaq glacier in a retracted position, c. 2 km from the little ice age maximum. aerial photographs from the early 1940s, 17 june 1948, 13 august 1968 and 20 july 1985 all show that the glacier front was located at a narrow part of the fjord. during this period, the glacier front had a surface falling steeply over a distance of c. 3 km from 500 m down to sea level. a landsat image from 2010 shows the position of the glacier front close to the position in 1968 (fig. 25). sermeq in sermilik icefjord the sermeq glacier is 65–70 km long and 3.5–5 km wide. on its way to sermilik icefjord the glacier receives several tributaries from local ice caps in the surrounding alpine highland (fig. 27). a river plain and tidal flat, c. 5 km long, separate the glacier front from the fjord (fig. 28). the present position of the glacier front is 63°32´n, 50°45́ w. the middle and upper part of the outlet is seen in fig. 29. the glacier is surrounded by very fresh moraines around its front, so the little ice age maximum is close to the present front position. about 6 km east of the present front, the glacier sends a branch northwards towards alanngorlia fjord. this branch is also separated from the fjord by a 1–2 km long river plain/tidal flat. the sources from before the middle of the 1800s provide little information about the position of the front of the sermeq glacier. after investigations of the fjord sermilik and its northern branch alanngorlia in 1878, jensen wrote (1879, pp. 31–32): the margin of the inland ice has lately advanced in these fjords and according to information by training college teacher kleinschmidt, boats could, as late as at the beginning of this century, pass through the valley that is found east of the mountain iviangiusat, but this strait is now completely filled with the inland ice. (authors’ translation.) if this is true, the front may have advanced at least 7 km since the beginning of the 1800s, supposing that the mentioned passage through alanngorlia first was blocked by proglacial sedimentation before the glacier passed and blocked this 4444 a b fig. 26 1878 1936 fig. 26. nakkaasorsuaq in bjørnesund. a: the watercolour of the glacier front was painted by a. kornerup on 15 june 1878 (kornerup 1879, 1978). b: the glacier front photographed by j. helk on 26 july 1936. copyright arktisk institut, copenhagen. j. helk was head of the photogrammetrical section at the geodetic institute in copenhagen and subsequently director of arktisk institut in copenhagen. 45 2 km iviangiusarsuit sermeq 63°30´n iviangiusarsuit alanngorlia 50°30´w 50°30´w50°30´w 63°30´n fig. 27 1985 fig. 27. sermeq glacier, aerial photograph from 20 july 1985, geodetic institute, route 886k no. 1843. dotted line: deformation of the median moraine according to a satellite image from 2010. meltwater from the glacier drains west (left) to the sermilik fjord. passage. today, the place name ‘iviangiusat’ is changed to iviangiusarsuit. with respect to early maps of the area, we note that the alanngorlia–sermilik connection was blocked by ice on maps by kleinschmidt from 1855 and by b. peters from 1859 (b. peters in the kleinschmidt map collection at the royal library in copenhagen (weidick 1959, fig. 42)). the first map by jensen (1889) was corrected in later versions (presumably after data acquired during later expeditions to west greenland) so that the glacier front of sermeq is shown in a position closer to the present one and with a distance to the alanngorlia branch of c. 6–8 km. the northern part of the front of sermeq is located at a ‘bay’ just south of a large local glacier coming from iviangiusarsuit (fig. 28). the following information on the position of the glacier front was given by k. gripp (1975) from investigations in 1930. gripp noted that moraines were formed at the active ice margin. a photograph of the glacier front shows the entire front and its surroundings, seen from the north at an elevation of 385 m a.s.l. the estimated position of the front in 1930 is shown on an aerial photograph from 1968 (fig. 28b). if it is correct that the position of the glacier front at the beginning of the 1800s allowed boats to pass from alanngorlia to sermilik fjord south-west around iviangiusarsuit, the front of the sermeq glacier may have had a position at least 5.5 km behind the position in 1930 and c. 6 km behind that in 1968 (figs 27, 28). the following advance, given by the estimated positions of the glacier front in 1985 and 2010, is shown in fig. 28. the estimated advance from c. 1800 to 1930 was c. 6 km (?), corresponding to 46 m per year. from 1930 to 1968 it advanced 18 m per year (700 m in 38 years) and from 1968 to 1985 (figs 27, 28b), it advanced 12 m per year (200 m in 17 years). finally, from 1985 to 2010 (landsat image) the northern flank of the glacier front may have advanced c. 100 m, corresponding to c. 4 m per year. isortuarsuup sermia isortuarsuup sermia at 63°50´n, 50°00´w, is an over 20 km long, c. 5 km wide outlet that calves in the isortuarsuup tasia lake at c. 450 m a.s.l. (fig. 29). the calf-ice 4646 63°33´n 1 km 1985 lia max ea rly 1 80 0s ? 1930? 2010 1968 sermeq iviangiusarsuit 63°33´ 50°37´w 50°37´w b a 1930 1968 fig. 28. a: the front of sermeq glacier seen from the north-west from an altitude of 385 m a.s.l. photograph by k. gripp, 11 september 1930, published by gripp (1975, plate 1, fig. 2). according to gripp, most of the drainage from the ice front appeared to occur via a subglacial tunnel located below the large median moraine at the centre line of the glacier front. the median moraine can also be seen on the aerial photograph of the front from 1968, shown in fig. 28b. b: sermeq glacier. aerial photograph from 14 august 1968, geodetic institute, route 281h, no. 263. the estimated positions of the glacier front in 1930 (see fig. 28a), 1985 and 2010 are shown on the photograph. lia max: little ice age maximum. 47 fig. 29. aster image of isortuarsuup sermia from 4 august, 2004 covering a large part of sermeq (fig. 27) and the un-named glacier 1cg14004 (fig. 31). the area adjacent to the margin of the inland ice is barren, partly alpine with local glaciers confluent with the marginal part of the inland ice, which might influence the response character of the ice-sheet margin. the position of the glacier front is close to the positions in 1985 and 2010. between 1938 and 1987 the ice-dammed lake at 710 m a.s.l. has drained at regular intervals (every 8th or 9th year) via isortuarsuup sermia. this periodicity is thought to have also prevailed since 1987 with inferred draining events in 1993 and 2002 (see fig. 30). 1cg14004 ice-dammed lake at 710 m a.s.l. isortuarsuup tasia 63°30´n 50°w 50°w 63°22´n 5 km isortuarsuup sermia 64°n se rm eq ser meq fig. 29 2004 64°n 50°w 2004 4848 production appears to be small and the maximum height of the glacier front is about 40 m above the lake level. details of the glacier and its surrounding area are best seen on an aerial photograph from 1968 (geodetic institute, route 281o, no. 317, 13 august). the glacier front was surrounded by a 100–300 m wide trimline zone indicating recent thinning of about 50 m near the front. the trimline zone pinches out at elevations above 750 m a.s.l. the trimline zone indicates a little ice age maximum about 400 m more westerly than seen in 1968. comparisons with aerial photographs from 1937(?) and 1985 reveal only a slight change in the position of the front and the same is seen on a landsat image from 2010. ‘lake 710 m’ is a large ice-dammed lake situated on the north side of isortuarsuup sermia, 5–10 km behind (east of) the front. the glacier and the ice-dammed lake are situated in a region that received little attention before the 1930s. due to the need for hydropower for nuuk (kangerluarsunnguaq or buksefjorden hydropower plant), hydrological and glaciological investigations have been carried out in order to calculate the future hydropower potential (braithwaite 1989). as illustrated by fig. 29, the margin of the inland ice around isortuarsuup sermia is surrounded by alpine uplands and highlands covered with large local glaciers. the determination of the position of the main glacier front must therefore to some degree be controlled by the mass balance and dynamics of the contributing local glaciers. the position of the front of isortuarsuup sermia varied little from 1950 to 2004. the stability of the frontal position and the thickness of the glacier are also reflected in the regularity of the draining periods of the ice-dammed lake at 710 m. between 1938 and 1987, draining occurred at intervals of 8–9 years (thomsen et al. 1992; fig. 30). the aster image from 2004 (fig. 29) shows a lake in the process of being filled up and it is possible that the lake drained in 1992 and 2001. in 2011, the lake level was again at a maximum (fig. 30). 1cg14004 1cg14004 is a small, c .6 km long and c. 1 km wide, unnamed land-based outlet from the inland ice, situated at 63°58´n, 49°42´w (figs 29, 31). the present glacier front is situated c. 1500 m behind moraines that mark the little ice age maximum. the age of the moraines is unknown, but we suggest that they formed in the midor late part of the 1800s, based on dating of the little ice age maximum of other smaller land-based glaciers in the region. the first information about the frontal position of 1cg14004 comes from aerial photographs from the 1930s (not very detailed) and from 1968 and 1985. it appears that in 1968 the front was situated c. 900 m behind the little ice age maximum. the recession continued up to 2010 when the glacier front was situated about 1500 m behind the little ice age maximum. kangaasarsuup sermia kangaasarsuup sermia is a land-based outlet with a length over 20 km from its beginning at the inland ice to the front at 64°06́ n, 49°59´w. at the front and along its flanks, the glacier is surrounded by a wide trimline zone, with a length of 2–3 km at the front (fig. 32). 647 m 695 m 653 m 747 m 729 m 750 750 800 800 751 m 1km ice margin maximum lake level minimum lake level 49°50´w 53°54.6´n 710 m b a 1940 1950 1960 1970 1980 1990 2000 2010 year 650 700 w at er le ve l ( m a .s. l.) fig. 30 fig. 30. ice-dammed lake at 710 m a.s.l. (see fig. 29) that is periodically drained via isortuarsuup sermia. the data for the period 1937–1987 are from thomsen et al. (1992), and the continuation is based on satellite images. a: estimated lake levels and draining events. b: map of the ice-dammed lake at 710 m a.s.l. redrawn from thomsen et al. (1992, figs 5.3 and 5.4, pp. 27 and 29). 49 historical knowledge about this outlet is sparse. the glacier is not shown on the first map of the area by j.a.d. jensen. subsequent investigations by the end of the 1800s formed the foundation for the map by amdrup et al. (1921; 1:1 000 000). descriptions of norse ruin groups no. 54 (nipaatsoq, c. 6 km west of the present glacier front) and no. 55 (kangaarsarsuk, at the north-western end of lake isortuarsuk, 3.5 km south of the glacier front) do not include any information concerning the glacier (jensen 1889, p. 117; bruun 1917, p. 102). the first sketch map showing details of the glacier is probably that published by roussell (1941, p. 14, fig. 6). a set of aerial photographs from 1936 (which formed the basis for roussell’s sketch map) shows the glacier front surrounded by a trimline zone (e.g. geodetic institute, route 768c, no. 25297) and on the aerial photographs the glacier front is situated c. 1.3 km behind the little ice age maximum as defined by the outer limit of the trimline zone. a later aerial photograph from 13 august 1968 (fig. 32) shows the front situated c. 3 km from the little ice age maximum. roussell (1941, p. 16) mentioned that according to thorhallesen, the eastern arm of ameralik (possibly naajat kuuat, fig. 10, loc. 9) reached a lake that was covered by an advancing glacier. it is unlikely that this lake was isortuarsuk (fig. 32), and it is also unlikely that the glacier was kangaasarsuup sermia. we consider it more likely that the lake was isvand, which was covered by the western flank of kangiata nunaata sermia in the 1700s (fig. 9, see earlier description). e. knuth (1944, pp. 94–109) provided a review of the knowledge of the topography of the interior parts of the nuup kangerlua region, based on his own and roussell’s archaeological investigations in the early 1930s, as well as on aerial photographs from 1936. knuth’s discussion included the ruin group ‘kangârssârssuk’. according to knuth, local people told him that the latter place name should be ‘kangârssârssuaq’ (knuth 1944, pp. 104–109). knuth included an aerial photograph of kangaasarsuup sermia (p. 99) and provided a short description of the frontal parts of the glacier, which he called a ‘dead glacier’ (p. 104) after crossing it. landsat and aster images from 2009 and 2010 record a continuation of the recession, with the front of the glacier situated c. 3.4 km behind the little ice age maximum in 2010. the trimline zone looks fresh and sterile, but possible correlation with the development around isvand north of the kangaasarsuup sermia may imply a greater age (1700?) for the formation of the outermost part of the trimline zone. north of kangersuneq to saqqap sermersua a wide coastal area with lowlands and uplands and numerous lakes is found between kangilinnguata sermia in the south and saqqap sermersua in the north (fig. 2). the ice-free coastland reaches elevations of around 1200 m a.s.l. near the margin of the inland ice. the c. 50 km long margin of the inland ice is undulating, with a number of lobes, but without proper outlets from the ice sheet. some parts of the ice margin terminate in ice63°56´n63°56´n 63°56´n63°56´n 1cg14004 49°36´w lia max 1 km 2010 1968 fig. 31 1968 fig. 31. aerial photograph from 1968 of outlet glacier 1cg14004 from the inland ice (geodetic institute, route 281s, no. 79, 13 august). the glacier code number is according to weidick et al. (1992). the approximate glacier extent during the little ice age maximum (lia max) and the estimated position of the glacier front in 2010 (from a satellite image) are shown. 5050 dammed lakes that drain to the south to the ujarassuit paavat and nuup kangerlua fjords. north of saqqap sermersua, ice-dammed lakes drain towards the søndre isortoq fjord (fig. 2). the ice-sheet margin around 65°n (fig. 2) is described below, an area for which some data are available from early visitors. specific descriptions focus on the isukasia area (sector 1ch02002; weidick et al. 1992), located in the central part of the region where the local hydropower potential has been investigated, and on the large outlet saqqap sermersua in the north. 1 km 64°05´n lia max lia max 2010 1936? 1948 64°05´n 50°w 50°w 1968 kangaarsarsuk (ruin group 55) kangaasarsuup sermia isortuarsuk naajat kuuat akuliarusersuaq kangaasarsuaq fig. 32 1968 fig. 32. aerial photograph from 13 august 1968 (geodetic institute, route 281q, no. 178) of kangaasarsuup sermia. the approximate positions of the glacier front during the little ice age maximum (lia max), in 1936 (geodetic institute, route 768c, no. 25297), in 1948 (geodetic institute, route 505 d1-ø, no. 9774, 4 september) and in 2010 (from a satellite image) are shown. the total recession is c. 3.4 km. 51 the ice-sheet margin at isukasia the southern areas of the ice-sheet margin at isukasia (65°11́ n, 49°50´w; fig. 33) were first mapped by j.a.d. jensen during 1884–1885 (jensen 1889). jensen’s map may be partly based on information from older sources. it is not possible on the basis of the outline of the lakes around isua and ataneq (fig. 34) to make comparisons with modern maps of the area, or to map changes in the position of the ice margin. the locality isua on jensen’s map is named isukasia on later maps (fig. 34). the area has been visited and described in connection with investigations of basement rocks and iron-ore deposits at isukasia (henriksen et al. 2009). glaciological investigations have been conducted as part of an evaluation of the hydropower potential of the locality in connection with possible future mining (colbeck 1974; kryolitselskabet øresund 1980). the glaciological investigations included determination of ice-margin fluctuations, going back to the 1930s, based on aerial photographs. during the period of glacial and hydrological investigations, direct measurements of annual fluctuations of the ice margin were carried out from august 1974 to august 1980. the data showed a continuous advance of the ice margin of 50–80 m, i.e. a few metres per year (kryolitselskabet øresund 1980). on the basis of aerial photographs and landsat images from 1969, 1985 and 2010 (fig. 33), the total advance between 1969 and 2009 is estimated to c. 300 m, or c. 7 m per year. the rate of advance seems to have been fairly constant during the last four decades, following the annual measurements from 1974 to 1980 by kryolitselskabet øresund (1980). saqqap sermersua saqqap sermersua at 65°12.5́ n, 50°39´w is a c. 28 km long, land-based outlet from the ice sheet. the frontal width is c. 3 km and at present a c. 2 km long alluvial plain separates the glacier front from the large tasersuaq lake (fig. 34). although remote, the saqqap sermersua area has attracted reindeer hunters. kleinschmidt’s map of the nuuk area (fig. 11) gives a good impression of the general knowledge of this area in the 1850s. the first published information on changes of saqqap sermersua appeared in the greenlandic newspaper atuagagdliutit (barselaj 1969 49°50´w 2010 1ch02002 isukasia 49°50´w 65°11´n 65°11´n lake at 792 m a.s.l. a 5 km 65°11´n fig. 33 1969fig. 33. part of the inland ice margin at isukasia. estimated advances of the ice margin near the lake at 792 m a.s.l. are shown on the aerial photograph from 1969 (geodetic institute, route 281y, no. 88, 10 august). the locality marked a, in the eastern inlet, is the site of annual measurements of frontal positions from 1974 to 1980 made by acg (kryolitselskabet øresund 1980). the approximate position of the glacier front position in 2010 is indicated. 5252 50°30´w 5 km 2009 65°15´n saq qap ser mersu a 1985 c.1850? 1936 1968 d c 65°15´n 50°30´w b a b d a c 65°n 65°n 50°w 50°w 20 km fig. 34 1985 fig. 34. aerial photograph of saqqap sermersua. estimated advance of the ice margin between the 1850s and 2009 is shown on the photograph from 11 july 1985 (geodetic institute, route 886k, no. 1456). the inset map was presumably made on the basis of the map by j.a.d. jensen after his survey work in the area in 1885 (amdrup et al. 1921). the locations of features a to d can be found on modern maps and the distances between these points on old and modern maps fit within ± 1 km; the date of this version of the map is unknown. the ice-front position around 1850–1860 is taken from fig. 11 (kleinschmidt’s map), and is located c. 2–3 km downstream of point d. 53 1866). barselaj was a parish clerk and a member of the local council in nuuk. his family name was probably ezekiassen, but only his first name was used as the author of the article. barselaj visited the saqqap sermersua area in 1862 and 1865. on his first visit he was told that the glacier was advancing. this was confirmed during his second visit to the glacier in july 1865, when he saw an old trail that was partly buried under the ice. the advance of the glacier appeared to be faster during the summer than in winter (h.c. petersen, personal communication, 1996). barselaj also noted that there was calf ice in lake tasersuaq, but the location of the front was not given, which is strange because at present the glacier front is separated from tasersuaq by an extensive alluvial plain. this plain is not shown on j.a.d. jensen’s map, made after surveys in 1884–1885. jensen’s map provides a relatively detailed picture of the area surrounding the glacier front (jensen 1889). however, there is no map scale and the coastlines and other landscape features are somewhat distorted. it appears that the front of saqqap sermersua was located close to the modern position. the map was used for decades as a base map of the area, and it was used by bruun to map norse settlements (fig. 7). a modified version of the map was included in the description of greenland, published 200 years after hans egede’s arrival in greenland (amdrup et al. 1921). part of this map with saqqap sermersua and the surrounding area (still without the alluvial plain in front of saqqap sermersua) is reproduced here and compared to an aerial photograph from 1985 (fig. 34). the suggested position of the glacier front shown on the aerial photograph in fig. 34 shows the probable situation in the 1850s, based on kleinschmidt’s map (fig. 11). north of the glacier, several large lakes are found (b). the lakes drain via a river (d) to saqqap sermersua at a place that was located 1 –2 km behind the front on kleinschmidt’s map. on aerial photographs (fig. 34) and modern maps, the glacier front is seen with the same river located 5–6 km behind the glacier front. this implies that the glacier front advanced c. 4 km from the middle of the 1800s to 1968 (c. 34 m per year). during the following period from 1968 to l985 it advanced c. 500 m (c. 29 m per year) and finally during the period from 1985 to 2009 it advanced c. 500 m (c. 21 m per year). we suggest that the front experienced a steady advance of c. 30 m per year throughout the period from the 1850s to the present time. the variations may reflect uncertainties in the determination of the frontal positions. discussion temperature changes considering the temperature history since the end of the 1700s in western greenland (vinther et al. 2006), the sparse information up to the mid-1800s may indicate a cold period; for the subsequent years (cappelen 2005), there is a parallel trend for all stations on the west coast with relatively cold periods around 1900 and a less extreme cold period c. 1990, interspersed with warmer spells c. 1940. the present marked increase in temperatures began c. 1995. the definition of geological periods in the quaternary is basically related to temperatures inferred from proxy data obtained from investigations of plant or animal remains or stable isotopes of material from sediment or ice cores. with respect to the term neoglacial, this period is defined by the cooling following the holocene thermal maximum that may have peaked in the inland areas of southern west greenland between 7000 and 6500 cal. years bp (bennike et al. 2010). the neoglacial can be subdivided in various ways (e.g. dahl-jensen et al. 1998; kaplan et al. 2002; seidenkrantz et al. 2007; kuijpers et al. 2009). it is subdivided here into an older (6500–2000 cal. years bp) and younger part (2000 cal. years bp to the present). older neoglacial (6500–2000 years bp) we can only provide a fragmentary picture of the fluctuations of the ice margin during older periods. the ice 5454 margin receded rapidly in the early holocene until the peak of the holocene thermal maximum at c. 7.0–6.5 cal. ka bp (fig. 35). at several localities reworked marine fossils have been found in neoglacial deposits or in glacier ice near the ice margin: 1. southern flank of the qassimiut lobe at around 61°n, 47°w. eight samples gave ages of 8.4 to 2.9 cal. ka bp (weidick et al. 2004). 2. søndre qoornoq bræ (61°09´n, 47°50´w. a few shells were found during reconnaissance work in 1955 (weidick 1959). no samples have been dated. 3. frederikshåb isblink (c. 62°37´n, 50°08´w). marine shells from the moraine and alluvial plain in front of the glacier were collected by d. heling (heling 1974) and a sample dated to 21 740 ± 400 14c years bp (weidick 1975a; i-7622). the sample is a bulk sample and may consist of a mixture of holocene and pre-holocene shells. 4. kangilinnguata sermia (c. 64°50´n, 50°00´w). four age determinations of shells gave ages from 6.4 to 4.3 cal. ka bp (table 2). 5. the jakobshavn isbræ area (around 68°58´– 69°45́ n, 50°14́ –50°20´w). fifteen radiocarbon dates of marine shells and a walrus tusk collected at the present margin of the ice sheet around jakobshavn isbræ ranged from 6.1 to 2.2 cal. ka bp (weidick & bennike 2007). in the area around jakobshavn isbræ, briner et al. (2010) have investigated numerous lakes close to the recent ice margin. they presented 18 radiocarbon ages ranging from 7.4 to 0.8 cal. ka bp. deglaciation continued after 7.3 cal. ka bp beyond the present ice margin until c. 2.3 cal. ka bp south of jakobshavn isbræ and until 0.4 cal. ka bp north of the ice stream. marine sediments at the mouth of jakobshavn isfjord (kangia) indicate ‘extensive ice phases’ at c. 2 and 0.5 cal. ka bp relating to a buried ‘narssarssuaq stade’ and a 1700–1800 culmination of little ice age advances in this region. the problem here is whether ‘extensive ice phases’ imply glacier advances or conversely extensive calf-ice production due to ice-front disintegration accompanying recession of a calving glacier. the general decrease in temperatures caused initial glacier expansions. beschel (1961) in his research on glacier fluctuations in west greenland discerned the general glacier advances in this period: the ‘hochmoos’ advance of an uncertain age, although indicated at c. 4000 years bp on his fig. 2, and the ‘larstig’ advance, estimated at between 2800 and 2500 years bp. both stades were defined in the alps from moraines found in front of little ice age moraines associated with minor local glaciers. the narssarssuaq stade in south greenland (61°10´n, 45°25́ w; weidick 1963; plate 1) is referred to the older neoglacial. the moraines of the narssarssuaq stade in the area around narsarsuaq in south greenland follow the present ice margin at a distance of 5–10 km along an ice-margin length of c. 60 km. a northern continu10 20 30 50 100 200 300 400 600 800 1000 d ep th (m ) 1930s warm period 1100 1200 1300 1330 1360 –40 δ18o (‰) –35 –30 10 20 50 100 200 500 1000 2000 5000 10 000 20 000 50 000 100 000 ag e (y ea rs b ef or e 19 58 ) little ice age medieval warm period holocene thermal maximum last glacial stage fig. 35 older neoglacial younger neoglacial roman warm period narssarssuaq stade fig. 35. late quaternary temperature history according to the camp century ice-core record. redrawn from dansgaard (2004). blue areas: colder than present; red areas: warmer than present. the age-depth model is tentative. 55 ation of this moraine system may have existed around the qajuuttap sermia ice stream, c. 30 km north-west of the narsarsuaq airport. moraines in this area were described by moltke & jessen (1896, p. 100). qajuuttap sermia started to advance and expand in the first half of the 1900s, an advance that probably ended towards the end of the last century. the moraines recorded by moltke and jessen were thus presumably removed or covered by qajuuttap sermia during this most recent advance (weidick 2009). the narssarssuaq moraines have been dated by lichenometry and radiocarbon dating to c. 2000 cal. years bp (dawson 1983; bennike & sparrenbom 2007). farther north at the head of kangerlussuaq (søndre strømfjord, 67°11́ n, 50°10´w, fig. 2), the age of a neo glacial advance of isunnguata sermia of the same magnitude as the following little ice age maximum was determined by optically stimulated luminescence (osl) dating to c. 2000 cal. years bp (forman et al. 2007). this advance may thus have taken place at the same time as the narssarssuaq stade. finally, a series of moraines on the nuussuaq peninsula north of disko bugt crosses the inner part of nuussuaq. the moraines can be followed from torsukattak icefjord in disko bugt in the south to ikerasaap sullua (qarajaq isfjord) in the north and continue on the peninsulas north of ikerasaap sullua. these moraines were first described by e. von drygalski (1897, vol. 1, p. 121) and later by weidick (1968, p. 92). the length of the moraine system across nuussuaq peninsula is 32 km and the system follows the present ice margin at a distance of 0.5–1 km. this glacial stade was called the drygalski stage by weidick (1968, p. 136). the moraines were possibly formed during a period when the relative sea level was not higher than at present, and the moraines are older than the little ice age moraines that are found between the drygalski moraines and the present ice margin. the moraine system is tentatively correlated to the narssarssuaq stade in south greenland. during reconnaissance work in august 1961, the system was followed all the way from torsukattak to ikerasaap sullua (beschel & weidick 1973, p. 313). younger neoglacial (2000 years bp – present) the medieval warm period (mwp) extended from ad 850 to 1200. it can be divided into several phases separated by one or more cold spells, as seen in records from ameralik fjord (kuijpers et al. 2009). farther south, kaplan et al. (2002) investigated a sediment core from ‘qipisarqo lake’, located 2 km from nordre qipisarqo bræ which is an outlet from the qassimiut lobe (fig. 2) at c. 61°03́ n, 47°42´w. between c. 9.1 and 0.4 cal. ka bp and again from ad 1850 onwards, the margin of nordre qipisarqo bræ was at or behind its present position. the qassimiut lobe had in general a reduced extent during the period 8.4–2.9 cal. ka bp, as shown by dating of reworked marine shells near its present margin (weidick et al. 2004). during the little ice age, nordre qipisarqo bræ was larger than at present, and it advanced into the catchment of ‘qipisarqo lake’, which led to deposition of minerogenic sediment in the lake during the period from ad c. 1550 to 1850 (kaplan et al. 2002). with respect to the advance during the little ice age, holst visited the margin of nordre qipisarqo bræ in 1880 (holst 1886; weidick 1959) and reported that the glacier front was situated only c. 100 feet behind the little ice age maximum. we suggest that the ice-margin position reached its maximum during the middle of the 1800s. a relatively cold period at around ad 600–900 is seen in the greenland ice-core records, such as the camp century ice core from 77°10´n, 61°08´w (dansgaard 2004; fig. 35). the record from the dye-3 ice core has been used for modelling the response of the inland ice margin north of jakobshavn isbræ, at paakitsup ilorlia (fig. 2; reeh 1983; weidick & bennike 2007). according to this model, a glacier advance occurred around ad 800; the magnitude of the advance was nearly as marked as later advances occurring around 1700 and 1900. there were possibly several cold spells, of which the most extreme may have been the so-called fimbul winter, described in nordic mythology, and dated to ad c. 536 (gräslund 2007). the initial onset of the little ice age is often placed at ad 1100, but the real start of the little ice age probably did not take place until after 1200–1400. in addition, there were warm spells between 1570 and 1600. these relatively warm periods subdivide the little ice age into several parts. no documented glacier advances took place during the oldest part of the little ice age although it is possible that traces of such advances were covered by younger advances in the 1700s and 1800s. according to the modelling by reeh (1983) of ice-margin fluctuations at paakitsup ilorlia near jakobshavn isbræ for the period ad 600–2000, an advance occurred around ad 1200, but the expansion of the ice margin was extremely modest. the trimline zone was mainly 5656 1800 1850 1950ad 1900 2000 1800 1850 1950ad 1900 2000 1800 1850 19501900 2000 2 4 6 0 2 4 km 0 2 km 0 2 km 0 2 km 2 0 1 km 0 2 km 0 1 km 0 2 km 4 6 km 8 10 12 14 16 18 20 22 km ? ? ? ? 1920 stade 1920 stade drainage of isvand 2004 saqqap sermersua ice margin west of isukasia kangilinnguata sermia narsap sermia qamanaarsuup sermia kangaasarsuup sermia 1cg14004 sermeq nakkaasorsuaq frederikshåb isblink (southern part) kangiata nunaata sermia ? last lia max? ?? 0 1 km isortuarsuup sermia last lia max? last lia max last lia max last lia max? ? advance lia max in mid 1800s? presumably close to lia max since 1700s last lia max 0 2 4 6 km ? ? separation of akullersuup sermia from kangiata nunaata sermia c. 1950 ? ? ? last lia max isvand area 0 frederikshåb isblink nuuk saqqap sermersua fig. 36 57 formed during the little ice age maximum from the 1700s or at the end of the 1800s. these two maximum advances are often marked by differences in vegetation cover and soil formation on the moraines, the older moraines being darker because they have a denser cover of lichens than the younger moraines. this was described for the qassimiut lobe, at qalerallit sermia in south greenland (weidick 1963, pp. 91–93; fig. 2). in some places, the old moraines (1700s?) are located distal to the younger moraines, whereas in other areas, the older moraines disappear under the younger ones. this contradicts the general idea, taken from local glaciers, that the oldest moraines are nearly always situated distal to the younger moraines. for small, local glaciers in west greenland, beschel (1961, pp. 1058–1059) followed this concept when dating moraines by lichenometry. an advance occurred at c. 1600 followed by another advance at 1770–1780. larger advances occurred in the second half of the 1800s, around 1870–1880 and 1890–1895, followed by a minor advance in 1920–1925. for the inland ice margin, the trimline zone often indicates the maximum advance from the 1700s, and the middle or late part of 1800s. however, in some sectors of the ice-sheet margin, the trimline zone is poorly developed, and in other places, a trimline zone is missing. information about fluctuations of the inland ice margin is sporadic. sometimes it is only based on lichenometric ages, which can be questionable (jochimsen 1973; webber & andrews 1973). however, lichenometric ages may be used to estimate the relative age of moraines, which are difficult to date using other methods. historical information can also be uncertain. an example discussed here is the uncertain determination of the position of the front of kangiata nunaata sermia, which in the middle of the 1800s, 1903 and in the l930s was at the approximately same position. this would lead to the conclusion that this glacier was relatively stable, with approximately the same position of the front for more than a century. however, the photograph by nissen from 1921 and the recognition of the ‘1920 stade’ show that the glacier front receded in the early 1800s, advanced c. 1903 and retreated in the 1930s (fig. 36). concluding remarks radiocarbon ages of marine material, transported from the subsurface of the west greenland ice sheet to its margin, demonstrate a widespread recession of the ice margin during the holocene. dated shells from the qassimiut lobe show that the ice margin was located behind the present margin as late as 2.9 ka bp (weidick et al. 2004) and at kangilinnguata sermia at 4.3 ka bp (this study). at the southern part of jakobshavn isbræ, the ice margin was behind the present margin at 2.2 ka bp (weidick & bennike 2007) and at the northern part of this ice stream at 0.4 ka bp (briner et al. 2010). these investigations focused on ice-free areas adjacent to the largest calf-ice-producing outlets from the ice sheet, and the behaviour of other parts of the ice-sheet margin may have been different. the fluctuations of the ice-sheet margin are still poorly documented, especially before the 1900s. however, from the sparse data available we can conclude that the inland ice margin had a retracted position, not only at the peak of the holocene thermal maximum, but also in the following millennia, at least in some lowland areas. local deviations may have occurred, such as the advance during the narssarssuaq stade (around 2000 years ago). this advance may have been related to the first large cold spell after the holocene thermal maximum, as shown in the temperature history inferred from the ice core from camp century (fig. 35; dansgaard 2004). a cold period facing page: fig. 36. estimated fluctuations of 12 major outlets and sectors of the inland ice margin discussed in the text. the curves illustrate the large variation in glacier response to climate change. three outlets show late little ice age maxima (lia max), in the 20th century, and other outlets have maintained an extent close to earlier little ice age maxima. only kangiata nunaata sermia shows large variations during the last two centuries. 5858 contours of ice surface (m) bedrock elevation (m) 1601–1700 1501–1600 1401–1500 1301–1400 1201–1300 1101–1200 1001–1100 901–1000 801–900 701–800 601–700 501–600 401–500 301–400 201–300 101–200 0–100 250 m interval 48°w51°w54°w 67°n 66°n 65°n 64°n 48°w51°w 63°n 62°n 1500 1000 1250 1750 2000 2250 2500 2750 12 50 22 50 20 00 17 50 15 00 67°n 66°n 65°n 64°n 63°n 62°n frederikshåb isblink s kns ns ss fig. 37 59 is recognised at 2000 years bp, which could have initiated the narssarssuaq stade that only locally went beyond the present extent of the ice margin. the subsequent cold period of the little ice age seems to have two minima at about ad 1550 and ad 1850 (dahl-jensen et al. 1998). traces of the early advances of the ice margin have been obliterated by later advances. with respect to the response of the individual sectors of the inland ice margin to climatic changes it must first be stressed that the current information about amplitude and exact dating of culmination is often uncertain, and that the trend of the fluctuations, even for the relatively well-documented last century shows a high degree of variability (fig. 36). it is clear that if the trimline zones around outlets indicate the maximum little ice age extent of the glacier, then localities without trimline zones must delineate areas where the present-day position coincides with the little ice age maximum. it is concluded that the little ice age maxima show a spread of ages, although for minor lowland outlets, the majority are related to the youngest little ice age maximum from the midor end of the 1800s. with respect to outlets without a trimline zone, it should be mentioned that some of these (narsap sermia, sermeq), at the end of the 1900s and the beginning of this millennium, seem to show an initial recession. in fig. 36 the fluctuations of the outlets and ice margin are shown. if this picture is compared with a map showing the topography of the ice-free coastland and the subglacial topography of the adjoining parts of the ice sheet, it can be suggested that the large fluctuations of kangiata nunaata sermia may be related to the presence of a large valley system that is evident under this ice stream (fig. 37), and therefore to the mass balance and the dynamics of this ice stream. otherwise outlets without trimline zones, such as saqqap sermersua in the north and isortuarsuup sermia in the south, seem to be related to nearby uplands and highlands although this can scarcely explain the steady advance of a local outlet. as regards the role of glaciers and their fluctuations as a ‘climatoscope’ for climate change in general, this seems to have been most successful for minor local glaciers where a well-defined delineation of the glacier form favours a more direct expression of the ruling local climatic and mass-balance conditions (leclercq 2012; leclercq et al. 2012). in contrast, the accumulation (catchment) area and ablation area within specific sectors of the inland ice may not be constant since the boundaries of these areas may shift position as a result of changing climatic conditions in combination with a series of factors that influence the dynamics and response time of the outlet. such factors include surface elevation, bed elevation, local and temporal variations in snow fall, mass balance, basal ice temperature, depth of the transition between weichselian and holocene age ice, curvature of the surface contours and slope aspect of the surface terrain (listed for modelling of the inland ice dynamics by ahlstrøm et al. 2008, p. 24). the specific gaps in understanding and modelling of the ice-margin response to climatic change are also dealt with by dahljensen et al. (2009); the observations and information discussed here can serve as examples of the strong variability in the response of the individual segments of the ice-sheet margin. this does not, however, weaken the credibility of estimates of the total mass balance of the entire inland ice. it is more related to the problem of prognoses of the response time of the individual sectors, where detailed short-term predictions are required to evaluate glacier hazards with respect to future exploration for minerals or hydropower. acknowledgements the authors are grateful to many people for help and support and for many discussions during the study. particularly to the following geus employees: andreas ahlstrøm, antoon kuijpers, camilla andersen, dirk van as, dorthe pedersen, horst machguth, karen edelvang, naja mikkelsen, robert fausto, signe bech andersen and søren nielsen. willy weng, frants van platen-hallermund and anette hindø helped with the facing page: fig. 37. map of southern west greenland showing the landscape in the coastal region and its continuation (subglacial topography) below the adjacent part of the inland ice. the marked fluctuations of kangiata nunaata sermia (kns) may be related to its role as the only lowland area that connects the marginal part of the ice sheet with the sea. the ‘quasi-stable’ behaviour of narsap sermia (ns) or the advance of sermeq (s) and saqqap sermersua (ss) may be related to interplays of mass balance, dynamics and the bedrock topography. data are from the etopo1 global relief model, provided by the national geophysical data center in the usa (amante & eakins 2009); suppressed colour tones are utilised to differentiate the subglacial topography from the surficial topography west of the inland ice. 6060 maps. dirk van as kindly provided information and photographs from recent field work at kangersuneq, and stuart watt helped to locate place names and gave advice on their spelling. henrik højmark thomsen and niels henriksen have supported us with knowledge on the earlier field work of the survey. h.c. petersen kindly provided information on the article by barselaj published in atuagagdliutit in 1866. rené forsberg discussed problems of determining recent uplift rates. the referees, ole humlum and niels t. knudsen, are thanked for their constructive and pertinent comments. data on the subsurface topography of the inland ice were provided by the etopo1 global relief model, provided by the national geophysical data center in the usa. aster data were distributed by the land processes distributed active archive center (lp daac), located at the u.s. geological survey (usgs) earth ressources observation and science (eros) center (lpdaac.usgs.gov). aerial photographs are published with the permission of kort & matrikelstyrelsen (a.200/87). dirk van as kindly provided photographs for the cover and the frontispiece. anker weidick dedicates this work to the memory of arne noe-nygaard, professor and head of the mineralogical museum in copenhagen from 1942 to 1978. inspired by studies by sigurdur thórarinsson on changes of icelandic glaciers from historical sources, noe-nygaard suggested that anker weidick should study fluctuations of greenlandic glaciers when he was a geology student in the early 1950s. 61 references ahlstrøm, a.p., mottram, r., nielsen, c., reeh, n. & andersen, s.b. 2008: evaluation of the future hydropower potential at paakitsoq, ilulissat, west greenland. danmarks og grønlands geologiske undersøgelse rapport 2008/37, 50 pp. alley, r.b., mayewski, p.a., sowers, t., stuiver, m., taylor, k.c. & clark, p.u. 1997: holocene climatic instability – a prominent wide spread event 8200 yr ago. geology 25, 483–486. amante, c. & eakins, b.w. 2009: etopo1 1 arc-minute global relief model: procedures, data sources and analysis. noaa technical memorandum nesdis ngdc-24, 19 pp. boulder, colorado: national geophysical data center. amdrup, g.c., bobé, l., jensen, a.s. & steensby, h.p. 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(by inclination i am greenlandic: about the life and work of kleinscmidt). det grønlandske selskabs skrifter 34, 528 pp. young, n.e., briner, j.p., axford, y., csatho, b., babonis, g.s., rood, d.h. & finkel, r.c. 2011: response of a marine-terminating greenland outlet glacier to abrupt cooling 8200 and 9300 years ago. geophysical research letters 38, l24701, http://dx.doi. org/10.1029/2011gl049639 http://dx.doi.org/10.1029/2005jd006810 http://dx.doi.org/10.1029/2011gl049639 http://dx.doi.org/10.1029/2011gl049639 65 akuliarusersuaq (akuliaruserssuaq), mountain north of kangaasarsuup sermia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 50 akullersuaq (akugdlerssuaq), semi-nunatak . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8, 10, 21, 24, 25, 26, 28 akullersuup sermia (akugdlerssûp semia), glacier . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7, 8, 9, 10, 14, 17, 22, 28, 29, 31, 56 alanngorlia (alángordlia), fjord . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 12, 43, 45 allumersat (agdlumersat), also bjørnesund, fjord . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7, 37, 42, 43, 44 ameralik, fjord . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8, 9, 10, 13, 15, 17, 20, 21, 23, 25, 49, 55 ameralla (ameragdla), fjord . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10, 11, 13, 17, 18, 21–23 amitsuarsuk, fjord . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8, 10, 11 anavik, see ujarassuit, ruin group . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7 13 arsuk bræ, glacier . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7 ataneq, river south-east of saqqap sermersua . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 51, 52 ‘auaitsirksarbik’ (giesecke 1910, pp. 257–258), not located . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19 austmannadalen (norwegian name, nansen 1890), valley . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8, 10, 11, 15–19, 21–24, 28 baals revier (bals revier), see nuup kangerlua, fjord . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17, 18 baffin bay . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6 bjørnesund, see allumersat . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 37, 44 buksefjorden, see kangerluarsunnguaq, fjord . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7, 9, 48 camp century, former station on the inland ice . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6, 54, 55 dalager nunatakker (danish name), nunataks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 39, 41 davis strait . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6, 38, 42 disko, island . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6, 12, 38 disko bugt, bay . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7, 11 dye-3, former station on the inland ice . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6, 55 eastern settlement (østerbygd), norse name for the settled area in south greenland . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 24 eqalorutsit kangilliit sermiat, now called qajuuttap sermia, glacier . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7, 55 eqalorutsit killiit sermiat (eqalorutsit kitdlît sermiat), glacier . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7 eqaluit, bay . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17, 18 frederikshåb isblink (danish name), glacier . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7, 37–43, 54, 56, 58 godthåb (danish name), see nuuk . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6, 7, 12, 13, 38, 56 index to place names this is an index to place names used in the text, together with the former spelling used in the original texts. many place names have been spelled in different ways over the years, in particular before 1851. after the greenlandic grammar was published by s. kleinschmidt (1851), the spelling of the greenlandic language was standardised, and kleinschmidt’s spelling was commonly used until 1973, when the present spelling was introduced. in this index, the place names show the 1973 spelling, followed by the kleinschmidt spelling, if different (in brackets). prekleinschmidt spellings and un-official names are in quotation marks. in addition, the feature relating to the names is given (mountain, glacier etc.). the exact location of old place names is often difficult to determine from old maps. the original spelling for stratigraphical and other geological terms is retained. thus the current spelling of a settlement is ‘kapisillit’, but the name for the local moraine stade is the kapisigdlit stade. similarly, the place name narsarsuaq is spelled so today, but we retain the spelling for the narssarssuaq stade. the location of ruin groups is seen in figs 3 and 7. the list does not show all the versions of spellings used by various authors. page numbers in bold refer to figures and tables. 6666 godthåbsfjord (danish name), see nuup kangerlua, fjord . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8, 13, 17, 18, 19, 37 gytjesø (danish name), lake . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8, 11 ikerasaap sullua (ikerasaup suvdlua), qarajaq isfjord, icefjord . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7, 55 illorsuit (igdlorssuit), ruin group 13a, on newer maps: illorssuakasiit (igdlorssuakasit) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8, 13, 24 indlandsisen (danish place name), inland ice, the greenland ice sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6, 7, 20, 21, 23, 59 isortuarsuk (isortuarssuk), lake . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8, 10, 17, 50 isortuarsuup sermia (isortuarssûp sermia), glacier . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7, 37, 45, 47, 48, 56, 59 isortuarsuup tasia (isortuarsûp tasia), lake . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 45, 47 isua, see isukasia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 51, 52 isukasia, area at the inland ice margin . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7, 9, 51, 56 isunnguata sermia (isúnguata sermia), glacier . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7, 55 isvand (norwegian name, nansen 1890), former ice-dammed lake . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7, 8, 10, 16, 17, 21–23, 28, 56 itilleq (itivdleq), passage or fjord . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10, 13 ivisaartoq (ivisârtoq), semi-nunatak . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 13, 34 iviangiusarsuit (iviangiussarssuit), earlier called ivianguisat, mountain . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 45, 46 j.a.d. jensen nunatakker (danish name), nunataks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 39, 40, 41, 42 jakobshavn isbræ (sermeq kujalleq), glacier . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7, 37, 38 54, 55, 57 jakobshavn isfjord (kangia), icefjord . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 54 johannes iversen sø (danish name), lake . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8, 11, 12, kangaarsarsuk (kangârsarssuk), ruin group 55 (bruun 1917, p. 102, roussell 1941, pp 13–14), called nûgasârsuaq by roussell . . 8, 13, 49 kangaasarsuaq (kangaussarssuaq), land area . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 50 kangaasarsuup sermia (kangaussarssûp sermia), glacier . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7, 8, 9, 10, 16, 17, 22, 28, 37, 49, 50, 56 kangaarsuup tasersua (kangârssûp tasersua), lake . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 39 43 kangerluarsunnguaq (kangerdluarssunguaq), buksefjorden (danish name), fjord . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 13, 48 kangerluarsunnguup tasersua, lake . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10 kangerlussuaq (kangerdlugssuaq), søndre strømfjord (danish name), fjord and airport . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6, 7, 38, 55 kangersuneq, fjord . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7, 8, 9, 10, 12–15, 18, 19–21, 23–26, 27, 29, 30–33, 36 kangia, norse settlement . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17 kangiata nunaata sermia (kangiata nunâta sermia), glacier . . . . . . . . . . . . . . . . . . . . . . 7, 8, 9, 10, 14–17, 21–29, 31, 37, 38, 49, 56–59 kangilinngua (kangilíngua), mountain . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 34, 36 kangilinnguata sermia (kangilínguata sermia), glacier . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7, 8, 9, 10, 12, 17, 34, 35, 36, 49, 54, 56 kapisillit (kapisigdlit), settlement . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8, 9, 10, 11, 15, 21, 26, 29 kapisillit kangerluat (kapisigdlit kangerdluat), fjord . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8, 10, 12, 18, 25, 26 karra (karra), mountain, lake . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8, 10 kilaarsarfik (kilârsarfik, roussell 1941: kilaussarfik), ruin group 51, ‘sandnes’ (norse name) . . . . . . . . . . . . . . . . . . . . . . 11, 13, 15, 17 kilunngaat (kilúngait), area at the inland ice margin . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 52 kuussuaq (kûgssuaq), river in austmannadalen, called ‘laxeelv’ by thorhallesen (p. 8.) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 18 ‘lake 8 m’, name used by iversen (1953) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 11 ‘lake 710 m’, ice-dammed lake . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 47, 48 langvand (norwegian name, nansen 1890), ice-dammed lake . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8, 16, 17, 21 22, 28 ‘laxeelv’ or laxelv (danish name), see kuussuaq, river in austmannadalen . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17, 18 majorariaq, river . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 39, 42 naajannguit (naujánguit), head of a fjord, bird cliff?,‘naviangoæt’ (bobé in thorhallesen 1914) . . . . . . . . . . . . . . . . . . . . . . . . . 34, 35 naajat kinginnerat (naujat kingingnerat), mountain . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 34, 36 naajat kuuat (naujat kûat), ‘storelv’ (danish name), river . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17, 18, 49, 50, nakkaasorsuaq (nákâssorssuaq), glacier . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7, 37, 38, 42, 43, 44, 56 nansens teltplads (danish name) (nansenip tupeqafia), nansen’s camp site . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8, 11 narsaq (narssaq), ruin group 12 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8, 24, 32, 33 narsap sermia (narssap sermia), glacier . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7, 8, 9, 10, 21, 24, 26, 32, 33, 34, 56, 58, 59 narsarsuaq (narssarssuaq), large plain, airport . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6, 7, 54 navianguat or naviangoæt, see naajannguit . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 32, 34, 35, 36 nikku (nivko), ‘nikok’ used by jensen 1889, mountain . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8, 19 67 nipaatsoq (nipaitsoq), ruin group 54 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8, 13, 49 nordre qipisarqo bræ, glacier . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7, 55 nordre sermilik, icefjord . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7 northgrip, ice core on the inland ice . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6 nunataarsuk (nunatârssuk), semi-nunatak . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8, 10, 21, 25, 26, 28 nunatarsuaq (nunatarssuaq), semi-nunatak . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8, 10, 21, 24, 28 nuuk (nûk), godthåb (danish name), capital of greenland . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6, 7, 12, 13, 17, 23, 38, 48, 53, 56 nuup kangerlua (nûp kangerdlua), godthåbsfjord (danish name), ‘baals revier’, fjord . . . . . . . . . . . . . . 6, 7, 9, 10, 13, 14, 15, 17, 29, 50 nuussuaq (nûgssuaq), peninsula . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7, 55 paakitsup ilorlia (pâkitsup ilordlia), fjord . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7, 55 paamiut (pâmiut), town . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 38 pisissarfik (pisigsarfik), mountain, sometimes also used for the nearby fjord (kapisillit kangerluat) . . . . . . . . . . . . . . . . . . . . 13, 20, 25 puilasoq (puilassoq), ruin group 8 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8, 13, 35, 37 qajuuttap sermia (qajûtap sermia), glacier, earlier eqalorutsit kangilliit sermiat . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7, 55 qalerallit sermia, outlet from the qassimiut lobe . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7, 57 qamanaarsuup sermia (qamanârssûp sermia), glacier . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7–10, 14, 17, 24, 25, 26, 28–32, 56 qaqqatsiaq (qáqatsiaq), mountain . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 43 qarajaq isfjord, see ikerasaap sullua . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7, 55 qassertup nuua (qassertup nûa), headland . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 29 qassimiut lobe, lobe from the inland ice . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7, 37, 38, 54, 57 ‘qipisarqo lake’, lake . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 55 qorlortoq (qordlortoq), water fall . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 39, 42 ‘sandnes’, see kilaarsarfik . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 11, 15 saqqannguaq (sarqánguaq), ruin group 11 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8, 13, 33 saqqap sermersua, glacier . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7, 38 saqqarsuaq (sarqarssuaq), ruin group 16 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8, 13, 14, 24, 25, 26, 28, 29 ‘saqqarsuaq lake’, suggested name for former ice-dammed lake . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25, 29, 31, 32 sermeq, also called ‘sermilik glacier’ (after the icefjord sermilik), glacier . . . . . . . . . . . . . . . . . . . 7, 12, 37, 38, 43, 45, 46, 47, 56, 58, 59 sermeq, situated in south greenland in søndre sermilik fjord, glacier . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7 sermeq kujalleq, jakobshavn isbræ (danish name), glacier . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 37 sermersuaq, ‘sermersoaq’ (giesecke 1910), used for an extensive glacier cover, varying from minor ice caps to the inland ice . . . . . . . . . 19 ‘sermilik glacier’, see sermeq . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7, 12, 37, 38, 43, 45, 46, 47, 56, 58, 59 sermilik, fjord . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7 sioraq, sandy beach, land area in front of frederikshåb isblink . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 39, 42 sisimiut, town . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6, 7, 15 summit, research station (the highest point on the inland ice) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6 søndre isortoq, fjord . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7 søndre qoornoq bræ (søndre qôrnoq bræ), glacier . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 54 søndre sermilik, fjord . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7 ‘storelv’, see naajat kuuat, river . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17, 18, 50 tasersuaq (taserssuaq), lake near saqqap sermersua . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8, 10, 52, 53 tasersuaq (taserssuaq), lake on north side of frederikshåb isblink . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 39 torsukattak (torssukátak), fjord . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7, 55 tummeralik (tungmeralik), ruin group 37 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8, 13, 19, 28 tummerallip tasersua (tungmeragdlip taserssua), lake . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8, 19 ujarassuit (ujaragssuit), anavik (norse name), ruin group 7 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8, 13, 35 ujarassuit (ujaragsluit), ujarachsoach, fjord . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17, 18, 29, 35 ujarassuit nunaat (ujaragssuit nunât), area . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 34 ujarassuit paavat (ujaragssuit pâvat), fjord . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8, 9, 10, 15, 17, 35, 36, 50 umiivik (umîvik), ruin group 15 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8, 12, 24, 28, 29 ‘umiviarsuit’, head of ameralla . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17 6868 vatnahverfi, norse site in south greenland . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 24 western settlement, vesterbygd (norse name), the settled area in the godthåbsfjord region in southern west greenland . . . . . . . . . 6, 24 1cg14004, un-named glacier . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7, 8, 10, 37, 47, 48, 49, 56 1cg14033, un-named glacier . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9 1ch02002, inland ice margin at isukasia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 51 _goback abstract introduction setting notes on the holocene history of the area geological, archaeological and historical information on glacier fluctuations and sea-level changes data from aerial photographs and satellite images glacier changes in the kangersuneq area the former ice-dammed lake isvand evaluation the kangiata nunaata sermia glacier system evaluation qamanaarsuup sermia evaluation narsap sermia evaluation kangilinnguata sermia evaluation comparisons with regional glacier fluctuations south of kangersuneq to frederikshåb isblink frederikshåb isblink nakkaasorsuaq sermeq in sermilik icefjord isortuarsuup sermia 1cg14004 kangaasarsuup sermia north of kangersuneq to saqqap sermersua the ice-sheet margin at isukasia saqqap sermersua discussion older neoglacial (6500–2000 years bp) younger neoglacial (2000 years bp – present) concluding remarks acknowledgements references index to place names research article | short dahl-jensen et al. 2021: geus bulletin 47. 5552. https://doi.org/10.34194/geusb.v47.5552 1 of 8 monitoring for seismological and geochemical groundwater effects of high-volume pumping of natural gas at the stenlille underground gas storage facility, denmark trine dahl-jensen1* , rasmus jakobsen1 , tina bundgaard bech1 , carsten møller nielsen1 , christian nyrop albers1 , peter h. voss1 , tine b. larsen1 1geological survey of denmark and greenland (geus), copenhagen, denmark abstract the large natural gas storage facility at stenlille, denmark, has been monitored to investigate the effect of pumping large amounts of gas into the subsurface. here, we present a new dataset of microseismicity at stenlille since 2018. we compare these data with methane in groundwater, which has been monitored since gas storage was established in 1989. further, we conducted a controlled 172 day microcosm experiment of methane oxidation on an isolated microbial community under both aerobic and anaerobic conditions. for this experiment, water was filtered from a well at stenlille with elevated levels of thermogenic methane and ethane. no microseismic activity was detected in the gas storage area above an estimated detection level of ml 0.0 for the established network. the long-term monitoring for methane in groundwater has still only detected one leak, in 1995, related to a technical problem during injection. the microcosm experiment revealed that oxidation of methane occurred only under aerobic conditions during the experiment, as compared to anaerobic conditions, even though the filtered water was anoxic. introduction storage of co2 in the subsurface as a means of reducing co2 in the atmosphere receives great international interest. therefore, there is a need for knowledge about how the subsurface behaves when large volumes of gas are pumped into reservoirs accompanied by potential contamination of groundwater aquifers. in denmark, the primary onshore interest focusses on sandstone reservoirs, with the gassum formation as a prime candidate (hamberg & nielsen 2000). the stenlille underground gas storage facility (fig. 1a) provides an opportunity to monitor the effects of large-volume pumping; while the gas pumped is natural gas and not co2, the volume pumped is large and can provide information about the effects of pumping activity on both groundwater geochemistry and microseismic activity. leakage is also important to monitor because groundwater aquifers are sensitive to changes (datry et al. 2004), and therefore increases in methane as well as trace concentrations of other alkanes may alter the groundwater ecosystem. groundwater aquifers are complex ecosystems that are of critical importance for geochemical cycles (griebler & lueders 2009). therefore, it has been important to monitor for hydrocarbons in the shallow groundwater *correspondence: tdj@geus.dk received: 08 sep 2020 accepted: 16 feb 2021 published: 22 mar 2021 keywords: co2 storage, carbon capture, geochemical monitoring, induced earthquakes, natural gas storage abbreviations: gc-fid: gas chromatography flame ionization detector ml: local magnitude mb: body wave magnitude geus bulletin is an open access, peerreviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: stefanie lode (geus) reviewed by: aaron cahill (heriot-watt university, uk) and one anonymous reviewer funding: see page 7 competing interests: none declared additional files: see page 7 https://doi.org/10.34194/geusb.v47.5552 https://orcid.org/000-002-0800-3105 https://orcid.org/0000-0003-1882-2961 https://orcid.org/0000-0001-8031-9326 https://orcid.org/0000-0002-1525-1385 https://orcid.org/0000-0001-7253-3509 https://orcid.org/0000-0002-6689-564x https://orcid.org/0000-0001-8240-337x mailto:tdj@geus.dk dahl-jensen et al. 2021: geus bulletin 47. 5552. https://doi.org/10.34194/geusb.v47.5552 2 of 8 www.geusbul let in.org because of the underground gas storage, and thereby detect possible impacts on this environment. here, we present a new dataset of microseismicity at stenlille since 2018 and compare these data with methane in groundwater, which has been monitored since gas storage was established in 1989. further, we report on a controlled 172 day microcosm experiment of methane oxidation on an isolated microbial community under both aerobic and anaerobic conditions. stenlille underground gas storage facility the stenlille underground gas storage facility is located 70 km south-west of copenhagen, operated by gas storage denmark a/s (fig. 1a). it was established in 1989 to buffer the supply of gas from the north sea and has since been re-developed to increase storage capacity. today, a total of 20 deep wells operate at the facility. fourteen wells are deployed for injection and withdrawal of gas and six wells are used for observational purposes. most observation wells are located in the periphery of the site (fig. 1b; laier & øbro 2009). the storage at stenlille is an anticlinal structure shaped by salt tectonics. a vertical closure of c. 35 m covers an area of 14 km2 (fig. 1c). the reservoir is formed in the upper triassic – lower jurassic gassum sandstone formation, where gas is stored by displacing formation water. the top gassum surface is located 1500–1600 m below ground level. the gassum formation consists of cyclically interbedded sandstone and marine mudstone deposited in a changing depositional environment (hamberg & nielsen 2000). the formation has excellent reservoir properties due to the overlying 300 m thick lower jurassic fjerritslev formation. it consists of claystone and hence serves as a caprock for the sandstone reservoir. the total estimated storage capacity of the stenlille structure equals three billion normal cubic metres (nm3), and due to reservoir heterogeneities, gas is stored in several separate zones. for safety and environmental reasons, the storage operation is monitored carefully. no sign of gas leakage has been observed in a monitoring well located in a sand stringer 15 m above the gas reservoir (laier & øbro 2009). other wells are monitored for possible lateral escape of natural gas. a baseline study of naturally occurring hydrocarbons (laier & øbro 2009) performed before the stenlille facility came into operation indicated the presence of only trace amounts hydrocarbon gases (in the form of biogenic methane) in the subsurface of the stenlille facility. the current amount of gas stored is just under 1600 mnm3, with an annual injection and extraction of close to 500 mnm3. the pumping rates typically vary between 100 and 250 knm3/h but with extraction rates occasionally going up to 400 knm3/h (fig. 1d). the detailed pumping activity for the individual wells for the period summer 2018 to summer 2019 can be seen in supplementary file s1. methods seismological monitoring during august and september 2018, we established a seismic network for monitoring microseismicity around the stenlille underground gas storage facility. the network consists of six seismographs placed within 5 km of the main pumping facility (fig. 1b). data for the period 1 october 2018 to 31 march 2020 have been screened for events, using the condet code (havskov et al. 2020). the screening triggers several hundred times on the data. the triggers are very unevenly spaced in time, depending not only on actual seismic events, but also on thunderstorms and noise. a manual screening of the triggered events resulted in 32 locatable seismic events and a large number of acoustic events related to thunderstorms. geochemical monitoring monitoring for hydrocarbons has been carried out monthly from 1989 to 1994, followed by quarterly measurements thereafter by the geological survey of denmark and greenland (geus). shallow groundwater is monitored from private drinking water wells, groundwater wells supplying waterworks in the vicinity of the gas storage and observation wells at stenlille. the two observation wells, k1 and k2, were established where the risk of leakage was considered highest. observation well k1 allows water samples to be taken from melt-water sand at 36 m depth and at 98 m in paleocene calcareous sand (fig. 2). observation well k2 was established in 1993 with a screen in the melt-water sand at a depth of 25–39 m (fig. 2). groundwater was sampled quarterly in 15 ml serum bottles with rubber septa and stored at 5°c until analysis within two days of sampling. analysis of c1–c4 hydrocarbons was done by gas chromatography flame ionization detection (gc-fid) on a shimadzu gc2010 equipped with a capillary column (gs-gaspro, 60 m, 0.32 mm) with a detection limit of 2 µg/l for both methane and larger hydrocarbons. microcosm experiment a total of 2400 l of groundwater was pumped and filtered over a glass fibre filter with a pore size of 0.3 µm and 293 mm in diameter (sterlitech, kent, washington, usa) using a submersible pump (grundfos mp1, grundfos, bjerrringbro, denmark). the pump rate was 8–10 l  min-1 (september 2019). the filter was stored https://doi.org/10.34194/geusb.v47.5552 http://www.geusbulletin.org dahl-jensen et al. 2021: geus bulletin 47. 5552. https://doi.org/10.34194/geusb.v47.5552 3 of 8 www.geusbul let in.org fig. 1 location, site details and pumping activity at the stenlille underground gas storage facility. a: the stenlille facility is located on sjælland, denmark (red hexagon) b: well locations and extension of the different gas zones at the stenlille underground gas storage facility. contour lines indicate depth in metres to the top gassum formation. red triangles indicate the stenlille seismic network seismographs deployed as part of the secure project. seismographs are within 5 km of the main pumping station (ste00). st-01 to -20 are wells and ste01–06 are seismic stations. modified from gas storage denmark a/s 2018 (fig. 2.2). c: the cross-section of the stenlille underground gas storage facility. prod.: production well. obs.: observation well. modified from gas storage denmark a/s 2018 (fig. 2.1). d: history of pumping activity at stenlille. modified from gas storage denmark a/s 2018 (fig. 3.1). 0 1 km n legend zone 1–3 gas wells zone 5 gas wells observation well (reservoir) observation well (cap rock) fault gas zone 1–3 (estimated nov. 2011) gas zone 5 (estimated nov. 2011) 15 00 15 00 14 90 1490 1480 14 7014 60 1470 1460 1450 1480 14 60 ste05 b ste00 ste01 ste02 ste06 ste03 st-06 st-01 st-05st-02 st-03 st-09 st-16 st-17 st-19 st-12 st-10 st-13 st-14 st-11 st-18 st-20 st-04 st-15 st-08 st-07 d 800 0 1600 t o ta l g a s [m io n m 3 ] gas extraction 1989 1991 1993 1995 1997 1999 2001 2003 2005 2007 2009 2011 2013 2015 2017 1570 clay, sealing d ep th ( m b e lo w s ea l l ev el ) d epth (m b e low seal level) 0 1550 1530 1510 1490 1470 1450 nw c se 1570 1550 1530 1510 1490 1470 1450 1 2 3 4 5 6 7 8 9 gas sand clay/silt, partly sealing obs. st-15 obs. st-4 prod. st-17 prod. st-11 prod. st-7 prod. st-1 storage zone 1 2b 3 4 5 6 obs. st-6 prod. st-19 7˚e 15˚e 54˚n 58˚n stenlille a https://doi.org/10.34194/geusb.v47.5552 http://www.geusbulletin.org dahl-jensen et al. 2021: geus bulletin 47. 5552. https://doi.org/10.34194/geusb.v47.5552 4 of 8 www.geusbul let in.org paleocene calcareous sand paleocene marine clay glacial till meltwater sand 500 m level (m) 60 40 20 0 –20 –40 –60 –80 k2 k1 wsw ene fig. 2 geological cross-section of the upper layers of the stenlille underground gas storage facility. approximate location and depth of monitoring screens in k1 and k2 are shown. the red shading indicates the probable distribution of gas after a leakage event at st-14 in august 1995. modified from the work of laier (2012). cold after sampling and microcosms were setup within 24 hours. the filter was divided into 24 fragments, and each part was placed in a 583 ml flask containing 90 ml of groundwater and 10 ml of salt solution. the concentrations in each microcosm were 0.62 mm mgso4•7h2o, 0.99 mm kno3, 0.88 mm nh4cl, 0.18 mm nahco3, 0.51 mm nacl, 0.52 mm cacl2•2h2o, 5.00 mm caco3 and 0.50 mm fe(iii)ooh. flasks were crimp sealed with butyl rubber stoppers and flushed with he before a controlled headspace was created comparing (1) untreated control, (2) ch4 and (3) ch4 and c2h6 under both aerobic and anaerobic conditions. flasks were incubated in the dark at 10°c on a rotational shaker at 75 rpm. during incubation, the concentrations of ch4, c2h6, o2 and co2 was quantified by gc-fid. results and discussion seismological events of the 32 located events, 20 were previously known (both earthquakes and explosions) and occurred close to stenlille (fig. 3) or were large enough and had a frequency content within the range used to trigger on the stenlille stations alone. twelve previously unknown and ‘spurious events’ (events of unknown origin) were found. these were only observed on the stenlille network along with many similar events, which could not be located. none of the events found are within the stenlille underground gas storage facility (fig. 3). the events identified are of a local magnitude (ml) of –0.2 to 2.5. a distant event with body wave magnitude (mb) 6.7 is observed and two regional events of ml 3.5 and ml 2.8 were also observed. the newly identified spurious events are all smaller than ml 1.0. a full list is provided in supplementary file s2. earthquakes and explosions these events, detected by screening the stenlille network, were all previously located by the routine monitoring of the danish seismic service. the fact that these events were also detected by independent screening of the stenlille seismic network confirms the fact that the network is capable of detecting such events (fig. 3). spurious events these events all have magnitudes smaller than ml 1 (fig. 3) – most are much smaller and all have similar signals, with strangely low frequencies in the range of 10 hz. this is much lower than expected for such small events. we do not know the cause of these events and their locations have an uncertainty of up to tens of kilometres. nonetheless, due to the difference in p and s wave arrival times (which relates to the distance from the station to the event; data not shown) we are convinced that they do not originate within the area of the stenlille underground gas storage facility. correlation with pumping we have detailed information on the pumping activity at the stenlille underground gas storage facility (see supplementary file s1). however, we observe no seismic events in the immediate vicinity of the stenlille facility, and so it is not possible to correlate events at stenlille with pumping activity. detection level in order to determine the detection level within the area covered by the stenlille facility we have used 10 natural earthquakes detected by the stenlille network. https://doi.org/10.34194/geusb.v47.5552 http://www.geusbulletin.org dahl-jensen et al. 2021: geus bulletin 47. 5552. https://doi.org/10.34194/geusb.v47.5552 5 of 8 www.geusbul let in.org the equation for calculating the ml for earthquakes in denmark is: ml = 0.925 × log10(a) + 1.61 log10(δ) – 2.38 (1) where a is the maximum amplitude of the s and surface wave arrival train in nm and δ is the distance in km (gregersen 1999). using the amplitude actually observed at each of the stenlillle stations for each of the 10 natural earthquakes and calculating what ml would be at distances covering the stenlille gas storage area, we obtain a series of estimates at which magnitudes it would be possible to see if the events occurred within the stenlille gas storage facility. the complete set of calculations are provided in supplementary file s3. average values for all earthquakes and stations are shown in table 1. we estimate the detection level of the stenlille seismic network to be at least ml 0.0. examples of noise analysis can be found in supplementary file s4. 8˚e 8˚e 9˚e 9˚e 10˚e 10˚e 11˚e 11˚e 12 e̊ 12˚e 13 e̊ 13˚e 14˚e 54˚n 54˚n 55˚n 55 n̊ 56˚n 56 n̊ 57˚n 57˚n 0 50 kmmagnitude 4 magnitude 3 magnitude 2 magnitude 1 fig. 3 seismological events detected by screening data from the stenlille seismic network. blue: earthquakes. light blue: explosions or presumed explosions. lilac: spurious events. red: earthquakes and explosions found by the danish seismological service in the period october 2018 to march 2020. triangles are seismological stations. dark green: the stenlille network. yellow: raspberry shake stations. light green: permanent stations in national networks. table 1 observed local magnitude (ml) around stenlille natural gas storage facility distance (km) 1 2 3 4 5 6 7 8 9 10 average observed ml (no units) –1.2 –0.7 –0.4 –0.2 0.0 0.1 0.2 0.3 0.4 0.4 https://doi.org/10.34194/geusb.v47.5552 http://www.geusbulletin.org dahl-jensen et al. 2021: geus bulletin 47. 5552. https://doi.org/10.34194/geusb.v47.5552 6 of 8 www.geusbul let in.org why are no events observed at stenlille? the stenlille network has, through 18 months of monitoring, not registered any seismic events within the stenlille underground gas storage facility. the detection level is estimated to be at or below ml 0.0, meaning that any missed events are very small. geus has monitored for earthquakes in denmark for many years, and no earthquake has ever been located at stenlille, bearing in mind that the detection level for the national monitoring network is higher at ml 2.0 (dahl-jensen et al. 2013; voss et al. 2015). the stenlille facility has been operational since 1989, pumping up to 500 mnm3 of gas in and out (fig. 1d) and has to our knowledge never received a complaint about shaking. we do not know the reason for the lack of events, but we can speculate. of course, events below our detection level could be occurring. it is possible that during the operation of the monitoring network, after almost 30 years of pumping, all stresses have long since been relieved. possibly events did occur early in the storage facility’s life. the stresses in the subsurface, both natural and established by the pumping itself, are too small to trigger earthquakes. geochemistry: preliminary results and longterm monitoring prior to operational onset the background concentration in the groundwater aquifers at stenlille contained only biogenic methane with concentrations <0.5 mg/l (laier & øbro 2009). hence, leaks from the gas storage site are simple to monitor due to low methane background concentrations and the lack of higher hydrocarbons. the injected gas consists of methane (89.5%), ethane (6.9%) and propane (2.6%; laier & øbro 2009); therefore, traces of ethane in groundwater would be a sensitive indicator of gas leakage. the analytical detection limit is c. 2 µg/l. during the 30 years of operation, there has only been one known leakage. in september 1995 there was a leak due to technical problems during gas injection, and even though it was quickly stopped, an estimated 5000 nm3 were lost to geological formations above the k1 m g c h 4/l 0 2 4 6 8 10 98–128 m 36–46 m k2 19 90 19 91 19 92 19 93 19 94 19 95 19 96 19 97 19 98 19 99 20 00 20 01 20 02 20 03 20 04 20 05 20 06 20 07 20 08 20 09 20 10 20 11 20 12 20 13 20 14 20 15 20 16 20 17 20 18 20 19 20 20 m g c h 4/l 0.0 0.1 0.2 0.3 0.4 0.5 25–39 m fig. 4 methane concentration in groundwater from monitoring wells k1 (91–128 m and 36–46 m depth) and k2 (25–39 m depth). in 1995, methane concentrations reached 27 mg/l in the deep k1 screen. https://doi.org/10.34194/geusb.v47.5552 http://www.geusbulletin.org dahl-jensen et al. 2021: geus bulletin 47. 5552. https://doi.org/10.34194/geusb.v47.5552 7 of 8 www.geusbul let in.org reservoir cap rock. for further details see the work of laier & øbro (2009). this leak has resulted in elevated thermogenic methane and ethane concentrations in the deep k1 monitoring with a decreasing trend since 1995 as shown in fig. 4. methane in groundwater aquifers can have different origins. thermogenic methane can rise from deeper or shallower hydrocarbon reservoirs into shallow sediment layers and aquifers due to natural gas migration, as observed from contamination in the deep k1 well. biogenic methane originates from methanogens that produce methane from acetate or hydrogen and co2 released when other microorganisms ferment organic matter in anoxic subsurface systems (beeman & suflita 1990; kleikemper et al. 2005). biogenic methane is detected in the k2 screen. during migration, methane can be oxidized by methanothrophic bacteria or archaea. this oxidation may occur aerobically in the presence of oxygen or anaerobically with nitrate, sulphate, and oxidized forms of iron and manganese. because the energetics of the anaerobic process are severely constrained, the process can take place only through syntrophic cooperation, involving interspecies electron transfer or other interdependencies. accordingly, anaerobic methane-oxidizing archaea organisms have never been obtained in pure cultures (knittel & boetius 2009). even though methane is often detected in groundwater with reducing redox conditions there is limited literature on methane oxidation in groundwater compared to marine and freshwater sediments. hence, from the thermogenic impacted k1 well, isolated microbial communities were established in controlled microcosms under both aerobic and anaerobic conditions to compare methane and ethane oxidation at stenlille. the concentration of methane and ethane were monitored for 172 days. during the incubation study the average methane oxidation rate was 35.97 µmol l-1 day-1 and 27.99 µmol l-1 day-1 in a headspace of methane or methane and ethane, respectively. the average ethane oxidation rate was 11.07 µmol l-1 day-1. oxidation of ch4 occurred only under aerobic conditions even though the isolated microbial community was adapted to anaerobic conditions from the k1 well. further, the added salt solution contained no3, so4 and fe(iii) that have been shown to work as an electron acceptor during anaerobic methane oxidation. lack of methane oxidation under anaerobic conditions has been observed elsewhere (cahill et al. 2017; kuloyo et al. 2020). implications for future co2 gas storage sites in denmark we report no seismic events at the stenlille facility during the monitoring period (october 2018 to may 2020) and no sign of elevated methane concentrations in the shallow groundwater that could be linked to operational activities. further, if small leaks are occurring from the underground storage there is a natural capacity in the microbial subsurface community to oxidize methane as observed in our microcosm experiments. based on our results, this oxidation will most likely take place in the transition zone between aerobic and anaerobic conditions, or potentially by an anaerobic community not captured in our experimental setup. it is important to detect leakage immediately. the seismic method described facilitates an online continuous detection and could be combined with additional geochemical sampling. this could supplement the quarterly monitoring, triggered by registered seismic events of a certain magnitude and location. this does not imply that seismic events necessarily lead to leakage, but it could be an additional tool to increase the likelihood of rapidly detecting a leak. acknowledgements the authors thank gas storage denmark a/s for hosting a seismic station and providing them with pumping data and methane samples. they are also thankful to the residents in the vicinity of the stenlille underground gas storage facility who kindly allowed them to place seismic monitoring stations on their properties. funding this work was carried out as part of the eu project secure, funded by the european union’s horizon 2020 research and innovation program under grant agreement number 764531. author contributions tdj: conceptualization, formal analysis, investigation, methodology, resources, visualization, writing – original draft, writing – review & editing. rj: resources, supervision, writing – review & editing. tib: data curation, methodology, visualization, writing – original draft, writing – review & editing. cmn: conceptualization, project administration, writing – original draft, writing – review & editing. cal: data curation, investigation, writing – original draft. pv: conceptualization, data curation, formal analysis, investigation, software, writing – review & editing. tbl: data curation, formal analysis, investigation, writing – review & editing additional files four additional files are available at https://doi.org/10.22008/fk2/t2251f references beeman, r.e. & suflita, j.m. 1990: environmental factors influencing methanogenesis in a shallow anoxic aquifer: a field and laboratory study. journal of industrial microbiology 5(1), 45–57. https://doi. org/10.1007/bf01569605 cahill, a.g., steelman, c.m., forde, o., kuloyo, o., ruff, s.e., mayer, b., mayer, k.u., strous, m., ryan, m.c. & cherry, j.a.j.n.g. 2017: mobility and persistence of methane in groundwater in a controlled-release field experiment. nature geoscience 10(4), 289–294. https://doi. org/10.1038/ngeo2919 dahl-jensen, t., voss, p.h., larsen, t.b. & gregersen, s. 2013: seismic activity in denmark: detection level and recent felt earthquakes. geological survey of denmark and greenland bulletin 28, 41–44. https:// doi.org/10.34194/geusb.v28.4717 datry, t., malard, f. & gibert, j. 2004: dynamics of solutes and dissolved oxygen in shallow urban groundwater below a stormwater https://doi.org/10.34194/geusb.v47.5552 http://www.geusbulletin.org https://doi.org/10.22008/fk2/t2251f https://doi.org/10.1007/bf01569605 https://doi.org/10.1007/bf01569605 https://doi.org/10.1038/ngeo2919 https://doi.org/10.1038/ngeo2919 https://doi.org/10.34194/geusb.v28.4717 https://doi.org/10.34194/geusb.v28.4717 dahl-jensen et al. 2021: geus bulletin 47. 5552. https://doi.org/10.34194/geusb.v47.5552 8 of 8 www.geusbul let in.org infiltration basin. science of the total environment 329(1–3), 215– 229. https://doi.org/10.1016/j.scitotenv.2004.02.022 gas storage denmark a/s. 2018: stenlille gaslager – undergrunden årsrapport 2017 rep. unpublished company report for gas storage denmark (in danish only). gregersen, s. 1999: national survey and cadastre (kms), copenhagen, denmark. in: emsc-csem newsletter no 15. european-mediterranean seismological centre. griebler, c. & lueders, t. 2009: microbial biodiversity in groundwater ecosystems. freshwater biology 54(4), 649–677. https://doi. org/10.1111/j.1365-2427.2008.02013.x hamberg, l. & nielsen, l.h. 2000: shingled, sharp-based shoreface sandstones: depositional response to stepwise forced regression in a shallow basin, upper triassic gassum formation, denmark. geological society, london, special publications 172, 69–89. https://doi. org/10.1144/gsl.sp.2000.172.01.04 havskov, j., voss, p.h. & ottemöller, l. 2020: seismological observatory software: 30 yr of seisan. seismological research letters 91(3), 1846–1852. https://doi.org/10.1785/0220190313 kleikemper, j., pombo, s.a., schroth, m.h., sigler, w.v., pesaro, m. & zeyer, j. 2005: activity and diversity of methanogens in a petroleum hydrocarbon-contaminated aquifer. applied and environmental microbiology 71(1), 149–158. https://doi.org/10.1128/ aem.71.1.149-158.2005 knittel, k. & boetius, a. 2009: anaerobic oxidation of methane: progress with an unknown process. annual review of microbiology 63(1), 311–334, https://doi.org/10.1146/annurev.micro.61.080706. 093130 kuloyo, o., ruff, s.e., cahill, a., connors, l., zorz, j.k., hrabe de angelis, i., nightingale, m., mayer, b. & strous, m. 2020: methane oxidation and methylotroph population dynamics in groundwater mesocosms. environmental microbiology 22(4), 1222–1237. https:// doi.org/10.1111/1462-2920.14929 laier, t. 2012: results of monitoring groundwater above the natural gas underground storage at stenlille, denmark. geological survey of denmark and greenland bulletin 26, 45–48. https://doi.org/10.34194/ geusb.v26.4748 laier, t. & øbro, h. 2009: environmental and safety monitoring of the natural gas underground storage at stenlille, denmark. geological society, london, special publications 313(1), 81–92. https://doi. org/10.1144/sp313.6 voss, p., dahl-jensen, t. & larsen, t.b. 2015: earthquake hazard in denmark. unpublished geological survey of denmark and greenland report, 2015/24, p. 53. copenhagen: geological survey of denmark and greenland. https://doi.org/10.34194/geusb.v47.5552 http://www.geusbulletin.org https://doi.org/10.1016/j.scitotenv.2004.02.022 https://doi.org/10.1111/j.1365-2427.2008.02013.x https://doi.org/10.1111/j.1365-2427.2008.02013.x https://doi.org/10.1144/gsl.sp.2000.172.01.04 https://doi.org/10.1144/gsl.sp.2000.172.01.04 https://doi.org/10.1785/0220190313 https://doi.org/10.1128/aem.71.1.149-158.2005 https://doi.org/10.1128/aem.71.1.149-158.2005 https://doi.org/10.1146/annurev.micro.61.080706.093130 https://doi.org/10.1111/1462-2920.14929 https://doi.org/10.1111/1462-2920.14929 https://doi.org/10.34194/geusb.v26.4748 https://doi.org/10.34194/geusb.v26.4748 https://doi.org/10.1144/sp313.6 https://doi.org/10.1144/sp313.6 monitoring for seismological and geochemical groundwater effects of high-volume pumping of natural g abstract introduction stenlille underground gas storage facility methods seismological monitoring geochemical monitoring microcosm experiment results and discussion seismological events earthquakes and explosions spurious events correlation with pumping detection level why are no events observed at stenlille? geochemistry: preliminary results and long-term monitoring implications for future co2 gas storage sites in denmark acknowledgements funding author contributions additional files references figures fig. 1 location, site details and pumping activity at the stenlille underground gas storage facility fig. 2 geological cross-section of the upper layers of the stenlille underground gas storage facilit fig. 3 seismological events detected by screening data from the stenlille seismic network. blue: ear fig. 4 methane concentration in groundwater from monitoring wells k1 (91-128 m and 36-46 m depth) an table table 1 observed local magnitude (ml) around stenlille natural gas storage facility geological survey of denmark and greenland bulletin 11, 61-86 61 origin and evolution of the kangâmiut mafic dyke swarm, west greenland kyle r. mayborn and charles e. lesher the kangâmiut dyke swarm in west greenland intruded archaean terrains at 2.04 ga, and its northern portion was subsequently metamorphosed to granulite facies during the nagssugtoqidian orogeny (c. 1.8 ga). mineral and whole-rock major and trace element compositions show that the parental magmas for the dyke swarm differentiated by the fractionation of olivine, clinopyroxene, plagioclase and late stage fe-ti oxides. petrographical observations and the enrichment of k 2 o during differentiation argue that hornblende was not an important fractionating phase. field observations suggest emplacement at crustal levels above the brittle–ductile transition, and clinopyroxene geothermobarometry constrains dyke emplacement depths to less than 10 km. granulite facies metamorphism of the kangâmiut dykes and their host rocks in the northern portion of the swarm requires subsequent burial to c. 30 km, related to roughly 20 km of crustal thickening between the time of dyke emplacement and peak metamorphism during the nagssugtoqidian orogeny. kangâmiut dykes are characterised by low ba/la ratios (12 ± 5), and high nb/la ratios (0.8 ± 0.2), compared to subduction related basalts (ba/la c. 25; nb/la c. 0.35). these geochemical characteristics argue that the kangâmiut dykes are not related to subduction processes. forward modelling of rare-earth element data requires that primitive magmas for the kangâmiut dykes originated from a moderately depleted mantle source with a mantle potential temperature of c. 1420°c. the inferred potential temperature is consistent with potential temperature estimates for ambient mantle at 2.0 ga derived from secular cooling models and continental freeboard constraints. the geochemistry and petrology of the kangâmiut dykes support a model that relates the dyke activity to passive rifting of the proposed kenorland supercontinent rather than to mantle plume activity or subduction. keywords: dyke swarm, laurentia, palaeoproterozoic, rifting _______________________________________________________________________________________________________________________________________________________________________ k.r.m., department of geology, western illinois university, macomb, il 61455, usa. e-mail: kr-mayborn@wiu.edu c.e.l., department of geolog y, university of california-davis, davis, ca 95616, usa. the 2.04 ga kangâmiut dyke swarm of west greenland has been the subject of numerous investigations over the past 35 years (windley 1970; escher et al. 1976; bridgwater et al. 1995; cadman et al. 1999), yet there is still considerable disagreement over many aspects of the swarm’s history. since the work of escher et al. (1976), who associated the dykes with synkinematic shearing during n–s compression, the swarm has often been cited as a type example of dykes that are emplaced into crustal regions undergoing shear deformation (cadman et al. 1999). the proposed compressional setting led cadman et al. (2001) to speculate that the swarm formed in a subduction-related environment. nevertheless, recent geochronology has shown that many of the shear zones originally thought to be consanguineous with the dykes are actually significantly older or younger than the dykes themselves (connelly & mengel 2000). the field area for escher et al.’s (1976) investigations is now known to contain archaean shear © geus, 2006. geological survey of denmark and greenland bulletin 11, 61–86. available at: www.geus.dk/publications/bull 62 zones and post-kangâmiut dyke shearing associated with the nagssugtoqidian orogeny. despite these complications, the kangâmiut dyke swarm offers a unique opportunity to constrain the magmatic and tectonic evolution of the province spanning the period from dyke emplacement to the nagssugtoqidian orogeny (2.04 to c. 1.8 ga). the present work discusses the constraints on the magmatic and metamorphic history of the kangâmiut dykes provided by field observations, petrology, geochemistry, and geochronology and critically evaluates previous and newly proposed models for their origin and subsequent metamorphism during the nagssugtoqidian orogeny. regional geology figure 1 is a geological map of central west greenland showing the nagssugtoqidian orogen and the kangâmiut dykes. the nagssugtoqidian orogen is bounded to the north by the palaeoproterozoic rinkian orogen (escher & pulvertaft 1976) and to the south by the archaean nain craton (nutman & bridgwater 1986; nutman & collerson 1991; friend & nutman 1994). the orogen is divided into four parts based on lithology, structure, stream sediment geochemistry, and aeromagnetic data. from north to south, these are the northern nagssugtoqidian orogen (nno), the central nagssugtoqidian orogen (cno), the southern nagssugtoqidian orogen (sno), and the southern nagssugtoqidian foreland (snf). these terrains are separated by three shear zones. the nordre strømfjord shear zone separates the nno from the cno, the ikertôq shear zone separates the cno from the sno, and the nagssugtoqidian front separates the sno from the snf. the nno contains archaean granitic gneisses and supracrustal rocks (van gool et al. 2002), whereas the cno is characterised by reworked archaean granitic and tonalitic gneisses, the 1.92 ga arfersiorfik quartz diorite, the 1.92 ga sisimiut charnockite, and supracrustal rocks (bak et al. 1975; kalsbeek et al. 1987; manatschal et al. 1998; kalsbeek & manatschal 1999; nutman et al. 1999; van gool et al. 1999). the sno and snf are composed of archaean granitic and tonalitic gneisses and the kangâmiut dykes. in the sno the dykes are metamorphosed to amphibolite facies in the south and granulite facies in the extreme north. the transition from amphibolite facies to granulite facies occurs within the ikertôq shear zone (korstgård 1979). in the snf, most of the kangâmiut dykes retain igneous textures and mineralogies. figure 1 shows that the ikertôq shear zone is the most continuous structure within the nagssugtoqidian orogenic belt and represents an important lithological boundary between the sno and the cno. it is traceable from the western shoreline, just south of the village of sisimiut, to the inland icecap. the shear zone contains panels of kangâmiut dyke-bearing tonalitic gneiss alternating with layers of garnet-sillimanite-kyanite-bearing metapelites, marbles, and quartzites that dip steeply nnw. the repetition of lithological units, the presence of down-dip lineations and s-vergent kinematic indicators show that this structure is a reverse fault (grocott 1979). deformed kangâmiut dykes are restricted to the southern footwall, while palaeoproterozoic felsic igneous rocks, including the 1.92 ga sisimiut charnockite and the 1.92 ga arfersiorfik quartz diorite, are found only north of the shear zone. u/pb dates for detrital zircons from the supracrustal rocks of the region require that these units were deposited after 2.10 ga (nutman et al. 1999), while 40ar/39ar dating of hornblende from metamorphosed kangâmiut dykes and orthogneisses from within the shear zone gives cooling ages of c. 1.73 ga (willigers et al. 1999). the itivdleq shear zone is located to the south of the ikertôq shear zone within the sno and is c. 6 km wide with e–w-trending lineations (hanmer et al. 1997). it contains multiple bands of sheared kangâmiut dykes and archaean orthogneisses that connelly & mengel (2000) used to document shearing during the archaean and the palaeoproterozoic nagssugtoqidian orogeny. for example, some of the shear bands are cut by weakly deformed tonalites that yield a magmatic age of 2498 ± 4 ma with a metamorphic overprint at 1782 ± 12 ma (connelly & mengel 2000). this shows that some of the shear bands near itilleq fjord developed prior to c. 2.5 ga and were subsequently intruded by tonalite. the metamorphic overprint age of 1782 ma and the presence of shear bands that cut the kangâmiut dykes indicate reactivation during the nagssugtoqidian orogeny. metamorphic hornblende from a recrystallised kangâmiut dyke within the itivdleq shear zone gives an 40ar/39ar age of 1873 ± 13 ma, consistent with the age of nagssugtoqidian metamorphism (willigers et al. 1999). the nagssugtoqidian front is the southernmost structural expression of the nagssugtoqidian orogenic event (fig. 1). hageskov (1995) showed that it is a discontinuous, en échelon array of nw-dipping, low-angle thrust facing page: fig. 1. geological map of the nagssugtoqidian orogen and its southern foreland in west greenland. nno, northern nagssugtoqidian orogen; cno, central nagssugtoqidian orogen; sno, southern nagssugtoqidian orogen; snf, southern nagssugtoqidian foreland. modified from van gool et al. (2002). 63 kangerlussuaq ikertooq northern cno ‘flat belt’ sisimiut nordre isortoqc n o sø nd re str øm fjo rd sø nd re str øm fjo rd sø nd re str øm fjo rd sukkertoppen iskappe inland ice attu n n o disko bugt qasigiannguit aasiaat itilleq kangaamiut maniitsoq sn o nordre strømfjord nordre strømfjord shear zone nordre iso rtoq shear z one ikertôq shear zone sn f nagssugtoqidian front kangerluarsussuaq 68° 67° 52° 50° 66° 50 km archaean palaeoproterozoic undifferentiated archaean gneisses aasivik terrane kangâmiut mafic dyke swarm supracrustal rocks sisimiut charnockite syntectonic granite suite arfersiorfik quartz diorite aasivik 64 faults that deformed the kangâmiut dykes in this region. the nagssugtoqidian front is well defined in the eastern portion of the orogen but is difficult to trace to the west approaching the sukkertoppen iskappe. it is not known whether the nagssugtoqidian front dies out beneath the icecap or swings to the nnw merging with the itivdleq shear zone (hageskov 1995). geological history the precambrian history of central west greenland involves five major events: (1) genesis and metamorphism of archaean crust; (2) emplacement of the kangâmiut dykes (2.04 ga); (3) deposition of sediments (c. 2.00–1.92 ga); (4) emplacement of the sisimiut and arfersiorfik intrusions (c. 1.92–1.87 ga); and (5) metamorphism and deformation accompanying the nagssugtoqidian orogeny (c. 1.82– 1.77 ga). genesis and metamorphism of archaean crust connelly & mengel (2000) and kalsbeek & nutman (1996) present u/pb dates and field observations that document the genesis and metamorphism of a large portion of the central west greenland crust during the late archaean. igneous zircons from foliated granulite facies gneisses give ages between 2.87 and 2.81 ga. granitoids that cut the foliation in these gneisses give ages between 2.81 and 2.72 ga. these cross-cutting relationships and dates indicate widespread genesis of granitic crust between 2.87 and 2.81 ga, immediately followed by metamorphism at 2.81–2.72 ga. emplacement of the kangâmiut dykes the emplacement of the kangâmiut dykes occurred after the formation of the archaean host rocks and before the nagssugtoqidian orogeny. u/pb geochronology on igneous zircons from three dykes gives ages of 2036 ± 5 ma, 2046 ± 8 ma (nutman et al. 1999), and 2048 ± 4 ma (connelly et al. 2000). the analysed zircons come from dioritic centres in wide composite dykes and cover most of the n–s extent of the dyke swarm. 40ar/39ar dating of hornblende from the kangâmiut dykes from the snf by willigers et al. (1999) gives emplacement related ages of 2.05–2.02 ga. deposition of sediments dating of detrital zircons from metasedimentary units just south of the nordre strømfjord shear zone and within the ikertôq shear zone (fig. 1) yields ages between 3.4 and 1.95 ga (nutman et al. 1999). some metasediments containing 1.95 ga zircons are cut by 1.92 ga quartzdiorites, requiring that deposition of the sediments occurred between 1.95 and 1.92 ga. metasediments found within thrust-bounded panels in the ikertôq shear zone contain zircons that yield ages between 2.1 and 2.0 ga. there are no known granitic intrusive rocks with ages between 2.5 and 1.92 ga within the central west greenland field area, suggesting that the sediments originated from a distal source. nutman et al. (1999) propose that the archaean and palaeoproterozoic terranes of eastern canada are possible source regions. sediments derived from these terranes would support the existence of the supercontinent kenorland (williams et al. 1991; aspler & chiarenzelli 1998). additionally, the deposition of sediments suggests the presence of basins that may have developed during the rifting and break-up of kenorland (van gool et al. 2002). the nagssugtoqidian orogeny and emplacement of sisimiut and arfersiorfik intrusions ramberg (1949) and noe-nygaard (1952) first recognised the nagssugtoqidian orogen based on the deformation and metamorphism of the kangâmiut dykes and the occurrence of shear zones with steeply dipping foliations. geochronological data show that the nagssugtoqidian orogenic event occurred between 1.91 and 1.77 ga (connelly et al. 2000) and resulted in granulite to amphibolite grade metamorphism and the development of discrete shear zones. pre-orogenic magmatism includes the emplacement of the 1.92 ga arfersiorfik quartz diorite (kalsbeek et al. 1987) and the 1.92 ga sisimiut charnockite. the orogenic event is proposed to be the result of continental collision that produced thrust stacking, folding and associated metamorphism (van gool et al. 2002). the suture between the two continents is not easily identifiable, but kalsbeek et al. (1987) proposed that the suture is located in what is now the boundary between the nno and the cno. 65 previous work on the origin of the kangâmiut dykes many early workers suggested that the emplacement of the kangâmiut dykes was directly related to the nagssugtoqidian orogeny. escher et al. (1976) stated that the dykes were emplaced into conjugate sets of active shear zones during nnw–sse compression. they based this hypothesis on their observations of conjugate sets of dykes that show a variety of cross-cutting relationships. they, and hanmer et al. (1997), proposed that the emplacement of both dyke sets occurred during shearing. bridgwater et al. (1995) expanded on the escher et al. (1976) hypothesis by proposing that the dykes formed during thrusting of amphibolite facies crust from the north under the granulite facies terrain in the southern nagssugtoqidian orogen. they invoked this hypothesis because the kangâmiut dykes contain hornblende suggesting they crystallised from a hydrous magma. bridgwater et al.’s (1995) hypothesis seeks to explain the hydrous nature of the kangâmiut dyke magmas by placing a hydrous source beneath the granulite facies host rocks. a potential problem with the escher et al. (1976) and bridgwater et al. (1995) hypotheses is that neither easily explains how partial melting of hydrous lower crust would directly produce melt of basaltic composition. if the escher et al. (1976) and bridgwater et al. (1995) hypotheses of dyke emplacement during compression are correct, then subduction related magmatism becomes a possibility. cadman et al. (2001) explicitly consider this possibility and proposed that the kangâmiut dyke swarm formed by adiabatic decompression of metasomatised mantle during passage of a slab window. the recent work of kalsbeek & manatschal (1999), connelly & mengel (2000) and van gool et al. (2002), offer alternatives to the subduction hypothesis and suggest that the swarm was emplaced during continental rifting. van gool et al. (2002) cite evidence of 2.0 ga rift related sediments in support for the rifting hypothesis. kalsbeek & manatschal (1999) speculate that the kangâmiut dykes are the product of mantle plume-related rifting based on the presence of ultramafic rocks found within the nagssugtoqidian orogeny. although these authors discuss the origin of the kangâmiut dykes, they do so only briefly, because their primary focus is on the nagssugtoqidian orogeny. field setting the kangâmiut dyke swarm intruded granulite facies archaean orthogneisses and is exposed over an 18 000 km2 area. figure 1 shows that the swarm extends for 150 km from just south of the village of maniitsoq towards sisimiut in the north and from the coast eastward to the ice cap. the dykes are most abundant near the coast and less so towards the ice cap. dyke widths range from a few centimetres to greater than 140 m. escher et al. (1975) estimated that dyke emplacement was accommodated by 2–3% crustal extension. appendix a provides locality and field characteristics of all the dykes examined in this study. field observations show that there are three dyke suites (mengel et al. 1996). two of these trend east–west, while the third has a ne trend. the ne-trending suite of dykes represents the vast majority of the dykes in the area and is the only suite that contains dykes with dioritic centres. the three sets of zircon populations used to date three separate dykes all come from these dioritic centres (nutman et al. 1999; connelly et al. 2000). thus, the netrending suite is dated at 2.04 ga and will be referred to as the ‘kangâmiut dykes’ proper as suggested by mengel et al. (1996). although the kangâmiut dykes have an overall ne trend, there are some systematic deviations. figure 1 shows that the southern portion of the swarm trends nne. moving northward, the orientation gradually changes to ene. changes in the orientation of the swarm correlate with increased dyke deformation and recrystallisation. escher et al. (1975) suggested that the bend in the swarm resulted from deformation during the nagssugtoqidian. hanmer et al. (1997) argue that this change in orientation is a primary feature related to the regional stress field during dyke emplacement. the majority of dykes south of the nagssugtoqidian front retain igneous mineralogies and textures with the exception of a small number of composite dykes with sheared dioritic centres. shearing was parallel to the dyke contacts, and mostly affected the large composite dykes. windley (1970) worked in an area just north of maniitsoq village where he observed cross-cutting dykes. he described a set of cross-cutting dykes where a younger dyke cuts the internal foliation of an older dyke. not all of the composite dykes show internal deformation. some dykes show irregular intrusive contacts between the dioritic centre and mafic host dyke showing that they formed by successive injections, closely spaced in time. large composite dykes show structural and petrological features not seen in the smaller non-composite dykes. for example, a 140 m wide dyke in kangerluarsussuaq fjord has fine-grained (c. 0.1–0.5 mm) equigranular mafic contacts. twelve metres from the contact the grain size is c. 1 mm with some contact-parallel primary layering. halfway towards the dyke centre the grain size increases 66 to c. 3 mm with primary clinopyroxene mostly replaced by hornblende. near its centre the dyke is slightly foliated and recrystallised to a garnet amphibolite. the centre of the dyke is a strongly foliated garnet-plagioclase-hornblende schist with a dioritic composition. a 10 cm wide epidote-calcite-quartz vein originates from the sheared centre and cuts the non-sheared mafic portion of the dyke. there are a variety of intrusive relationships between dykes and host rocks including en échelon steps, bridges, and forks. figure 2 shows bridges with sharp, angular edges contained within the chilled margin of a kangâmiut dyke. some dykes have chilled margins up to 40 cm thick, while others have no chilled margins. chilled margins contain fine-grained plagioclase and clinopyroxene phenocrysts in a microcrystalline groundmass. some chilled margins contain dismembered bridges of host rock, whereas dyke interiors contain no bridges or xenoliths. in the northern portion of the dyke swarm, metamorphic minerals and deformation features replace igneous minerals and primary intrusive features. in the itilleq fjord region, most of the dykes are partly to completely altered during static or shear-related recrystallisation. some dykes show penetrative foliation, whereas other dykes are deformed only at the contacts and form boudins within the deformed country rock. the central portions of these boudins are partially recrystallised. on the south shore of itilleq fjord, away from most of the nagssugtoqidian deformation, the dykes preserve primary emplacement structures, but are statically recrystallised to garnet amphibolites. dykes at the northern extent of the swarm within ikertooq fjord are completely recrystallised to granulite facies. petrography the chilled margins of the kangâmiut dykes contain 0.2– 0.8 mm phenocrysts of clinopyroxene and plagioclase in a microcrystalline groundmass of plagioclase, clinopyroxene, hornblende, quartz, and fe-ti oxides. plagioclase phenocrysts are euhedral to subhedral and weakly zoned, whereas the clinopyroxene phenocrysts are subhedral to anhedral. some chilled margins contain 0.2–1.5 mm hornblende crystals. figure 3a shows that these hornblende crystals contain abundant inclusions of fe-ti oxides, which are absent from the clinopyroxene and plagioclase phenocrysts. these hornblende crystals also enclose plagioclase and clinopyroxene phenocrysts suggesting they are a later phase that crystallised in situ. the interiors of most kangâmiut dykes are fineto medium-grained with subophitic textures. subhedral to anhedral clinopyroxene, plagioclase, and fe-ti oxides are the primary constituents along with interstitial quartz, hornblende, and trace amounts of apatite. in some of the dykes, anhedral clinopyroxene fills interstitial areas between subhedral to euhedral plagioclase. clinopyroxene displays two types of exsolution. the first type is laminar exsolution of low-ca pyroxene. the second type appears to be granular exsolution of low-ca pyroxene around the outer parts of the original clinopyroxene crystal. additionally, most clinopyroxene grains have rims of hornblende (fig. 3b). kangâmiut dyke samples from itilleq fjord are variably metamorphosed to fine-grained (0.2–1.0 mm) amphibolites with well-developed foliations. metamorphic assemblages include hornblende, plagioclase, quartz, garnet, titanite, and biotite. hornblende replaces clinopyroxene, whereas plagioclase and garnet form at contacts between hornblende and plagioclase (see also mengel et al. 1996). titanite replaces fe-ti oxides. metamorphic orthopyroxene occurs within kangâmiut dykes in ikertooq fjord near the northern extent of the swarm (fig. 1), marking the transition from amphibolite to granulite facies. the typical assemblage in these granulite facies dykes is plagioclase + hornblende + orthopyroxene + clinopyroxene ± garnet ± titanite (see also korstgård fig. 2. kangâmiut dyke with angular bridges within the chilled margin (photo: david bridgwater). 67 1979). these dykes are weakly foliated and fine-grained (0.1–1.0 mm). petrology and geochemistry whole-rock major and trace elements a total of 122 dyke samples were analysed for most major and minor elements on fused glass discs using a wavelength dispersive x-ray fluorescence (xrf) spectrometer at the geological survey of denmark and greenland (geus) in copenhagen. na 2 o was determined by atomic absorption spectrometry. kystol & larsen (1999) describe the analytical methods, precision, and accuracy of the geus lab and report that the standard error for all major and minor elements is less than 0.25 wt%, based on multiple analyses of international standards. trace element concentrations for 73 dyke samples were measured at the university of california-davis using a 1.0 mm plag hbl cpx a hbl plag cpx 0,5 mm b fig. 3. a: photomicrograph of a chilled margin of a kangâmiut dyke with clinopyroxene (cpx) and plagioclase (plag) phenocrysts in a groundmass of the same plus fe-ti oxides. the large crystal in the centre is hornblende (hbl) with inclusions of a plagioclase phenocryst and groundmass plagioclase and fe-ti oxides. sample ggu 430267, plane polarised light. b: photomicrograph of a sample from the interior of a kangâmiut dyke showing hornblende (hbl) rims on clinopyroxene (cpx). sample ggu 430999, plane polarised light. 68 perkin-elmer elan 500 inductively coupled plasma mass spectrometer (icp-ms). samples were prepared for analyses using the method described by jenner et al. (1990) with the exception that we utilised microwave digestion bombs to insure total dissolution. table 1 presents representative major and trace element data for the kangâmiut dykes, and figs 4 and 5 show these data in covariation diagrams. the full data set is available upon request from the first author. the kangâmiut dykes cover a range from 9.0–0.9 wt% mgo and the majority of the dykes would be classified as medium-k basalts or low-k basalts based on their k 2 o and sio 2 contents (le maitre 2002). first order observations of the major and compatible trace element data, as described below, indicate that the differentiation of the parental magma(s) of the kangâmiut dykes was influenced by the fractionation of plagioclase, clinopyroxene, late stage feti oxides and possibly olivine. first order observations based on covariation diagrams neither support nor refute the involvement of hornblende as a fractionating phase. figure 4c shows al2o3 concentrations that range between 11.4 and 15.8 wt%. the highest mgo sample has a low al2o3 concentration, whereas the next group of dykes at c. 7.0–7.5 wt% mgo have higher concentrations of al2o3. this increase of al2o3 between c. 9 and 7 wt% mgo likely reflects olivine and/or clinopyroxene fractionation. clustering of data between 7.5 and 4.5 wt% mgo defines a trend of decreasing al2o3 with decreasing mgo, indicative of plagioclase fractionation. the initial increase in total feo (feo + 0.9 × fe2o3) between c. 8.0 and 4.5 sample 430904 430923 430926 430931 430952 430970 430981 430988 430997 dyke 5 13 14 16 27 37 42 45 54 sio2 50.87 52.78 50.64 50.73 52.19 48.90 50.71 49.30 56.06 tio2 2.47 2.31 1.94 0.88 1.62 1.66 1.02 1.45 0.79 al2o3 12.62 13.65 13.29 12.57 13.19 13.14 14.09 13.34 21.46 fe2o3 4.27 3.53 4.26 1.74 4.14 2.52 1.55 1.91 1.24 feo 11.67 11.34 9.99 9.10 11.05 11.98 9.89 11.40 4.34 mno 0.24 0.20 0.22 0.21 0.24 0.24 0.21 0.22 0.08 mgo 4.27 3.51 5.55 8.90 4.57 6.40 7.43 6.60 0.92 cao 8.93 8.24 9.86 12.31 8.48 10.85 11.27 11.39 9.07 na2o 2.47 2.91 2.55 2.27 2.80 2.28 2.15 2.25 4.19 k2o 0.69 1.13 0.58 0.18 0.87 0.28 0.18 0.30 0.51 p2o5 0.26 0.29 0.19 0.07 0.19 0.13 0.09 0.13 0.16 loi 0.84 0.63 0.94 0.91 1.07 1.34 0.79 1.31 0.79 sum 99.60 100.52 100.03 99.88 100.40 99.72 99.39 99.58 99.61 sc 38 30 36 58 44 46 47 45 14 v 368 355 333 389 318 381 296 362 29 cr 47 14 63 145 43 63 153 98 5 ni 40 28 65 113 40 74 88 60 1 co 46 44 47 52 44 61 51 56 14 rb 22 33 16 3.9 25 5.6 4.3 6.1 10 sr 171 201 241 117 207 160 136 160 290 y 47 34 29 19 35 27 18 27 34 zr 182 168 139 47 133 99 46 92 157 nb 14.7 10.9 11.8 3.3 9.2 9.2 3.2 8.2 11.3 ba 204 321 156 50 306 90 50 95 192 la 17.5 20.5 17.1 4.4 12.7 9.2 4.4 8.4 12.3 ce 43.8 46.0 40.2 10.6 31.9 21.8 10.7 20.7 32.0 pr 6.07 6.13 5.41 1.49 4.43 3.10 1.70 3.00 4.54 nd 27.8 27.1 22.7 7.37 18.8 14.2 7.31 13.9 20.6 sm 7.14 6.11 5.44 2.45 4.58 3.93 2.32 3.71 5.09 eu 2.08 1.84 1.66 0.82 1.56 1.33 0.89 1.22 1.88 gd 7.52 6.84 5.55 3.01 5.42 4.51 2.76 4.16 5.41 tb 1.33 1.12 0.87 0.51 0.95 0.79 0.50 0.73 0.90 dy 8.17 6.80 5.22 3.12 5.63 4.61 3.10 4.37 5.44 ho 1.62 1.28 1.06 0.68 1.22 0.99 0.69 0.93 1.15 er 4.76 3.73 2.96 1.90 3.39 2.75 2.12 2.56 3.23 tm 0.70 0.52 0.42 0.30 0.52 0.42 0.29 0.38 0.48 yb 4.52 3.34 2.67 1.97 3.53 2.61 1.96 2.54 3.11 lu 0.63 0.49 0.39 0.29 0.54 0.42 0.29 0.37 0.45 hf 4.78 4.60 3.68 1.27 3.54 2.77 1.44 2.47 4.13 ta 1.03 0.76 0.82 0.26 0.64 0.54 0.24 0.52 0.72 pb 4.04 5.81 3.26 1.49 6.50 1.84 1.08 3.35 3.78 th 2.83 3.80 2.12 0.50 3.47 0.90 0.42 0.90 1.88 u 0.71 0.93 0.54 0.14 0.79 0.24 0.11 0.24 0.49 table 1. major and trace element data for representative kangâmiut dykes major element oxides in wt%; trace elements in ppm. sample numbers refer to geus databases. dyke localities shown in fig. a1 (appendix). 432102 432108 432115 432118 432122 432133 432138 432143 432158 86 60 64 64 64 71 70 76 75 57.18 48.97 51.01 56.84 50.25 51.00 49.46 50.57 50.48 1.65 1.69 1.92 2.12 1.82 1.64 2.70 1.33 1.34 14.37 13.75 12.97 13.07 13.67 13.96 15.57 13.14 13.59 2.15 4.73 2.21 1.98 2.17 2.06 2.54 1.26 1.75 8.94 9.48 12.73 10.99 12.22 11.59 11.51 11.99 11.84 0.16 0.23 0.24 0.19 0.23 0.21 0.20 0.23 0.23 2.34 6.12 5.08 2.14 5.81 4.78 3.44 6.98 6.12 6.11 10.72 9.46 6.17 10.12 9.53 8.61 11.40 10.59 3.69 2.44 2.57 3.37 2.45 2.46 2.99 2.11 2.07 1.14 0.39 0.52 1.43 0.33 0.47 1.02 0.22 0.35 0.35 0.15 0.19 0.39 0.10 0.18 0.44 0.10 0.13 1.43 1.41 1.01 1.17 0.91 1.64 1.35 1.11 1.06 99.50 100.08 99.90 99.86 100.07 99.53 99.83 100.44 99.55 20 38 41 22 41 36 29 49 45 155 348 393 207 558 303 259 390 336 33 134 83 22 94 61 33 174 103 17 82 54 15 66 49 23 78 74 28 53 46 32 54 49 32 53 52 31 7.4 13 45 8.0 12 28 4.6 8.5 306 235 164 247 164 200 213 138 142 35 26.2 37 48 23 30 53 24 28 194 106 133 326 78 110 200 58 84 18.1 9.3 10.0 23.9 6.2 7.3 15.8 4.6 5.2 392 104 145 454 98 156 366 69 106 27.8 10.86 12.9 42.0 7.4 11.4 18.2 5.8 7.6 63.9 26.28 30.6 94.4 18.1 26.7 40.4 14.1 18.0 8.66 3.73 4.36 11.96 2.58 3.83 5.58 2.07 2.67 36.2 16.93 19.1 49.7 11.6 17.3 25.7 9.90 12.2 8.12 4.35 5.05 10.85 3.02 4.71 6.91 2.72 3.24 2.49 1.45 1.68 2.93 1.10 1.47 2.20 1.02 1.20 6.98 4.46 5.27 10.02 3.51 4.86 7.39 3.27 3.99 1.11 0.75 0.99 1.47 0.60 0.82 1.32 0.57 0.68 6.31 4.47 6.23 8.26 3.75 5.01 8.56 3.68 4.51 1.20 0.98 1.24 1.60 0.81 1.08 1.88 0.82 1.04 3.19 2.44 3.71 4.22 2.28 2.92 5.03 2.36 2.89 0.45 0.34 0.57 0.61 0.34 0.41 0.70 0.36 0.41 2.89 2.31 3.45 3.90 2.20 2.67 4.97 2.31 2.81 0.44 0.35 0.55 0.55 0.32 0.43 0.74 0.35 0.41 4.68 2.83 3.49 7.64 2.10 2.78 5.19 1.59 2.36 1.27 0.64 0.66 1.55 0.41 0.50 0.98 0.30 0.34 3.65 1.92 2.63 4.89 1.73 3.07 6.66 1.03 1.88 4.52 1.19 1.73 7.92 1.00 1.61 3.88 0.60 1.04 1.10 0.29 0.46 1.92 0.28 0.40 0.87 0.16 0.29 69 wt% mgo is also indicative of plagioclase fractionation (fig. 4d). figure 4e shows that cao ranges from 12.4 to 6.0 wt% and correlates positively with mgo, indicating that clinopyroxene and/or plagioclase was part of the fractionating assemblage. figure 6 shows the cao/al2o3 ratio for the kangâmiut dykes decreases with decreasing mgo, and fig. 5a shows decreasing sc with increasing zr. both of these observations further indicate that clinopyroxene was a fractionating phase. the decrease in nickel with increasing zirconium, shown in fig. 5c, is related to clinopyroxene and/or olivine fractionation. figure 4b shows tio 2 concentrations that range from 0.9–3.5 wt% with trends that show increasing tio 2 from 9–4.5 wt% mgo that changes to decreasing tio 2 below c. 4.5 wt% mgo. figure 4d shows that a similar trend is a b c d e f g h 5 0 5 5 6 0 sio2 4 5 4 6 8 10 12 14 cao 0.5 1.5 2.5 3.5 tio2 10 12 14 16 18 al2o3 4 8 12 16 20 feo* 1.5 2.5 3.5 4.5 na2o 0 0.0 0.4 0.8 1.2 1.6 2.0 mgo k2o 2 4 6 8 10 0.0 0.2 0.4 0.6 mgo p2o5 0 2 4 6 8 10 fig. 4. variations of sio 2 , tio 2 , al 2 o 3 , feo, cao, na 2 o, k 2 o, p 2 o 5 with mgo (in wt%) for the kangâmiut dykes. all analyses are recalculated on an anhydrous basis with all iron as feo. 70 observed for total feo. additionally, vanadium (fig. 5d) shows an initial increase, then decrease with increasing zirconium. these changes from increasing to decreasing tio2, feo and vanadium concentrations suggest fe-ti oxides fractionated when the magmas reached c. 4.5 wt% mgo. mineral chemistry major element compositions of pyroxenes, plagioclase, and hornblende were acquired using a cameca sx-50 microprobe at the university of california-davis. analyses were made using a 15 kv accelerating voltage, a 10 na beam current, and a 1 µm beam. elements were calibrated using mineral standards. the data in tables 2–4 give the compositions of clinopyroxene, plagioclase, and hornblende phenocrysts from chilled margins. table 2 presents major and minor element compositions of clinopyroxene phenocrysts from chilled margins based on the average of 2–4 spot analyses of 4–7 grains per sample. the proportions of enstatite, ferrosilite, and wollastonite components are 0.48–0.52, 0.13–0.23, and 0.28–0.35, respectively. al2o3 concentrations range from 3.93–2.75 wt%, while na2o varies from 0.36–0.26 wt%. table 3 presents plagioclase phenocryst compositions from chilled margins. plagioclase phenocryst cores have an anorthite (an) component range of an69–an55. table 4 presents compositions of the hornblendes found in the chilled margins of some kangâmiut dykes. they would be classified as ferrohornblende and ferrotschermakite based on the classification of leake et al. (1997). whole-rock and mineral compositions in projection space figure 7 shows pseudo-ternary projections of whole-rock and mineral data for the kangâmiut dykes. the components cpx, plag, ol, and qtz were calculated using major and minor elements and the scheme of tormey et al. (1987). in the projection from qtz (fig. 7a), the dyke data form a cluster that is displaced from the centre 1 0 2 0 3 0 4 0 5 0 6 0 a b c d sc sr ni v zr zr 5 0 150 250 350 450 0 4 0 8 0 120 0 5 0 100 150 200 250 0 5 0 100 150 200 250 100 200 300 400 500 fig. 5. variation of sc, sr, ni, and v with zr (in ppm) for the kangâmiut dykes. 0 2 4 6 8 1 0 mgo 0.2 0.4 0.6 0.8 1.0 1.2 cao/al2o3 pl cpx ol fig. 6. variation of cao/al 2 o 3 with mgo. arrows show path caused by fractionation of olivine, clinopyroxene, or plagioclase from a starting composition with 9 wt% mgo and ca/al 2 o 3 = 1.0. 71 of the ternary plot towards the plag apex. in this projection all the dykes lie within the phase volumes defined by the joins between olivine, clinopyroxene and plagioclase (ol:cpx:pl) or hornblende, clinopyroxene and plagioclase (hbl:cpx:pl). in the projection from the cpx component (fig. 7b), the dykes define an array that projects away from both the pl:ol and the pl:hbl joins. similarly, in the projection from the plag component (fig. 7c), the dykes form an array that intersects both the cpx:ol and cpx:hbl joins. these observations show that plagioclase, clinopyroxene and either olivine or hornblende were cofractionating phases. a and c in fig. 7 also show experimentally determined 1 atm and 0.8 gpa ol:cpx:pl cotectics derived from melting experiments using a primitive kangâmiut dyke (dyke #45) as the starting material (mayborn 2000). the wholerock data form a cluster near the low pressure ol:cpx:pl cotectic in the projection from the qtz component. similarly, in the projection from the plag component the whole-rock data also plot close to the low-pressure cotectic. incompatible trace element behaviour figure 8 shows representative chondrite normalised rareearth element (ree) patterns for the kangâmiut dykes. the dykes are slightly light rare-earth element (lree) enriched with a la/smn ratio ranging from 1.16 to 2.50, with an average of 1.50. the heavy rare-earth (hree) patterns have shallow slopes with a dy/ybn range of 1.05 sio2 52.34 50.68 51.88 51.17 52.01 51.36 51.85 51.59 tio2 0.47 0.65 0.44 0.61 0.45 0.34 0.44 0.60 al2o3 3.17 3.93 3.14 3.21 3.34 2.75 3.63 2.68 cr2o3 0.11 0.19 0.18 0.21 0.22 0.08 0.18 0.07 feo 8.29 9.83 8.76 12.45 9.43 13.79 9.72 12.94 mgo 16.46 15.18 15.70 15.62 15.54 16.94 15.62 15.37 cao 19.59 19.18 18.96 16.16 19.01 14.32 18.01 16.52 na2o 0.29 0.32 0.26 0.29 0.28 0.36 0.31 0.26 sum 100.72 99.96 99.31 99.72 100.28 99.94 99.76 100.03 si 1.917 1.886 1.927 1.913 1.920 1.922 1.920 1.928 al(iv) 0.083 0.114 0.073 0.087 0.080 0.078 0.080 0.072 al(vi) 0.053 0.058 0.065 0.055 0.065 0.074 0.079 0.046 ti 0.013 0.018 0.012 0.017 0.012 0.011 0.012 0.017 cr 0.003 0.006 0.005 0.006 0.007 0.001 0.005 0.002 fe3+* 0.022 0.039 0.000 0.013 0.004 0.016 0.000 0.010 fe2+ 0.232 0.266 0.274 0.377 0.287 0.401 0.308 0.395 mn 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 mg 0.899 0.842 0.870 0.871 0.855 0.858 0.862 0.856 ca 0.768 0.765 0.755 0.647 0.752 0.611 0.715 0.661 na 0.021 0.025 0.020 0.021 0.020 0.036 0.022 0.019 k 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 total 4.011 4.020 4.001 4.006 4.002 4.008 3.996 4.005 mg# 0.78 0.73 0.76 0.69 0.75 0.67 0.74 0.68 wo 0.35 0.34 0.35 0.29 0.34 0.28 0.33 0.30 en 0.52 0.51 0.50 0.50 0.49 0.49 0.50 0.48 fs 0.13 0.16 0.16 0.22 0.17 0.23 0.18 0.22 sample 430988 432108 432145 432148 432158 430211 430267 430283 dyke 45 60 76 77 75 82 84 85 n = 7 n = 8 n = 6 n = 8 n = 8 n = 8 n = 10 n = 10 table 2. microprobe analyses of clinopyroxene phenocrysts from chilled margins oxides in wt%; n = number of grains. fe3+* calculated using the method of papike et al. (1974). dyke localities shown in fig. a1 (appendix). sample 430988 432108 432145 432148 432158 430211 dyke 45 60 76 77 75 82 n = 8 n = 8 n = 8 n = 8 n = 7 n = 8 sio2 52.18 49.74 53.36 53.79 51.23 54.13 al2o3 30.45 31.49 29.65 29.65 30.91 29.03 feo 0.65 0.57 0.59 0.74 0.96 0.92 cao 13.11 14.13 12.43 12.23 13.34 11.24 na2o 4.02 3.36 4.33 4.64 3.89 4.84 k2o 0.05 0.05 0.09 0.07 0.09 0.17 sum 100.46 99.34 100.45 101.12 100.41 100.34 ab 35 30 38 40 34 43 an 64 69 61 59 65 55 or 0.3 0.3 0.5 0.4 0.5 1.0 table 3. microprobe analyses of plagioclase phenocrysts from chilled margins oxides in wt%; n = number of grains. dyke localities shown in fig. a1 (appendix). sample 432158 430267 430283 dyke 75 84 85 n = 12 n = 14 n = 6 sio2 43.90 42.63 40.88 tio2 1.04 1.92 1.64 al2o3 10.69 10.05 9.98 feo 21.09 21.73 24.02 mno 0.22 0.19 0.21 mgo 8.09 8.05 6.44 cao 10.93 10.78 10.58 na2o 1.18 1.63 1.58 k2o 0.99 0.90 0.90 cl 0.57 0.67 0.65 f 0.09 0.05 0.14 sum 98.79 98.60 97.02 si 6.549 6.425 6.346 ti 0.117 0.218 0.192 al(iv) 1.451 1.575 1.654 al(vi) 0.427 0.209 0.172 fe3+ 0.767 0.799 0.922 fe2+ 1.864 1.940 2.196 mn 0.027 0.025 0.027 mg 1.798 1.809 1.491 ca 1.747 1.741 1.760 na(m4) 0.253 0.259 0.240 na(a) 0.089 0.217 0.237 k(a) 0.189 0.172 0.178 total 15.28 15.39 15.41 table 4. microprobe analyses of amphiboles from chilled margins oxides, cl and f in wt%. dyke localities shown in fig. a1 (appendix). 72 plag cpx ol proj. from qtz 0.8 gpa 1 atm 0.8 gpa 1 atm proj. from cpx ol plag 0.5 qtz 0.5 qtz cpx plag 0.5 qtz proj. from ol 0.5 qtz proj. from plag cpx ol 0.5 qtz 0.5 qtz ol:pl 17:83 hbl:pl 55:45 hbl:cpx:pl 40:20:40 ol:cpx:pl 10:40:50 kangâmiut dykes clinopyroxene hornblende plagioclase a b d c ol qtzplag +cpx ol qtz +plag cpx +ol qtzplag cpx fig. 7. pseudo-ternary projections of whole-rock and mineral compositions from the kangâmiut dykes using the projection scheme of tormey et al. (1987). dashed lines connect possible crystallisation assemblages. solid lines through data show the trend of kangâmiut dykes where this is well defined, and the locations of the olivine:clinopyroxene:plagioclase cotectics at 0.8 gpa and 1 atm. a: projection from the qtz component. b: projection from the cpx component. c: projection from the plag component. d: projection from the ol component. to 1.46 and an average of 1.15. the two most lreeenriched samples come from sheared dioritic centres and have dy/ybn ratios of 1.46 and 1.42, respectively. the patterns for the sheared dioritic centres cross-cut the patterns for non-sheared dykes. figure 9a shows incompatible trace elements for the kangâmiut dykes on primitive mantle normalised compatibility diagrams (thompson 1982). the elements on the right hand side of fig. 9a are moderately incompatible, with incompatibility increasing towards the left using the element order from sun & mcdonough (1989). the most primitive kangâmiut dykes have the lowest concentration of incompatible elements and have flat patterns with small negative nb and zr anomalies. the negative sr anomaly seen in the more evolved samples reflects plagioclase fractionation. the two most evolved samples 73 of peridotite is the mantle, then the basaltic nature of the kangâmiut dykes indicates that they resulted from partial melting of mantle peridotite followed by intracrustal differentiation. three settings where mantle melting occurs beneath continental crust are (1) subduction zones, (2) active rifts associated with mantle plumes, and (3) passive rift settings. each of these settings can have distinctive mantle compositions and conditions for melting 5 1 0 1 0 0 5 0 0 la ce pr nd sm eu gd tb dy ho er tm yb lu 430931 432143 430988 432115 432138 432102 432118 sample/chondrite 3 1 0 1 0 0 430931 432143 430988 432115 432138 432102 432118 rb ba th u nb la ce sr nd zr sm eu ti dy y yb lu sample/primitive mantle sample/primitive mantle 0 . 5 1 1 0 1 0 0 rb ba th u nb la ce sr nd zr sm eu ti dy y yb lu n-morb oib cfb island arcs primitive k-dykes continental arcs a b fig. 8. rare-earth element compositions of representative kangâmiut dykes normalised to c1 chondrite. chondrite normalising values from sun & mcdonough (1989). fig. 9. a: normalised incompatible trace element compositions of representative kangâmiut dykes. b: comparison of the kangâmiut dykes with normal mid-ocean ridge basalts (n-morb), ocean-island basalts (oib), island arc basalts, continental flood basalts (cfb), and continental arcs. representative ‘primitive kangâmiut dykes’ is the average composition of 18 kangâmiut dykes with a mgo range of 8.9–6.1 wt%. data for n-morb and oib are from sun & mcdonough (1989). island arc data compiled from bailey et al. (1989), pearce et al. (1995), and gust et al. (1997). cfb data compiled from hooper & hawkesworth (1993), lightfoot et al. (1993), wooden et al. (1993), peate & hawkesworth (1996) and storey et al. (1997). continental arc data compiled from tormey et al. (1991) and bacon et al. (1997). (432102 and 432118) are sheared dioritic centres and display negative ti anomalies. they also show a depletion of hree, as noted in fig. 8. figure 9b shows representative patterns for basalts from a variety of tectonic settings and a pattern representing the average of kangâmiut dykes with mgo > 6 wt%. notable features in the kangâmiut pattern are a smooth, slightly increasing trend from right to left through the moderately incompatible elements (lu to sm), small negative anomalies of nb, sr, and zr, and a relatively flat trend for th, ba and rb. the kangâmiut dyke pattern is distinct relative to the ocean island basalt (oib), midocean ridge basalt (morb), island arc basalt (iab), and continental arc basalt (cab) patterns. for example, the kangâmiut dykes do not display the enrichment of highly incompatible elements, or the positive nb anomaly seen in the oib pattern, or depletions seen in the morb pattern. relative to iab and cab, the kangâmiut dyke patterns do not exhibit the large negative nb anomaly or the positive sr anomaly. overall, the kangâmiut dyke pattern most closely resembles the pattern for continental flood basalts (cfb), including the shallow slope of the pattern, and the small negative nb and sr anomalies. discussion tectonic setting one of the primary goals of this study is to constrain the tectonic environment during emplacement of the kangâmiut dykes. experimental melting studies over the past 40 years have shown that basalts are products of partial melting of peridotite (reay & harris 1964; takahashi 1986; baker & stolper 1994). since the dominant source 74 that are reflected in the trace element compositions of the associated basalts. the kangâmiut dykes have evolved compositions with a mg-number (defined as 100mg/(mg + fe) on a molecular basis) range of 0.60–0.21. magmas in equilibrium with mantle peridotite will have a mg-number close to 0.71 (roeder & emslie 1970; langmuir et al. 1992) showing that even the most primitive kangâmiut dykes represent somewhat evolved magmas. thus, the major elements reflect both the fractionation and mantle melting histories. however, most incompatible trace element ratios remain relatively constant during crystallisation and can be used to constrain primary source characteristics. the following discussion examines some characteristics of the mantle source and the conditions of mantle melting revealed by examining incompatible trace elements from the more primitive dyke samples (with mgo > 4.5 wt%). subduction hypothesis for generation of kangâmiut dykes subduction environments generate basaltic melts by two different mechanisms of partial melting. the first is melting induced by lowering the solidus temperature of the peridotite by the introduction of volatiles from the subducting slab and subsequent decompression melting within the mantle wedge (jakes & gill 1970; tatsumi 1989; arculus 1994). the second is decompression melting associated with back-arc spreading (tatsumi et al. 1989; gribble et al. 1998). although the mechanisms of melting in these settings are different from each other, they both produce magmas with a compositional ‘subduction component’ that is indicative of a hydrated and metasomatised mantle. some important characteristics of subduction zone basalts are hfse depletions, lile enrichment, and high al 2 o 3 . the available data can be used to evaluate the subduction hypothesis for the kangâmiut dyke swarm implied by escher et al. (1976) and bridgwater et al. (1995), and explicitly proposed by cadman et al. (2001). as shown in fig. 9b, arc basalts have distinctive depletions in hfse. these hfse depletions occur in palaeozoic, proterozoic, and archaean arc-related basalts, suggesting that modern style subduction occurred in the archaean (stern et al. 1994; blichert-toft et al. 1995). in addition to hfse depletions, arc basalts are enriched in lile (e.g. pb, k, ba, rb, and cs) relative to basalts from other tectonic settings. this enrichment is proposed to 0 1 0 2 0 3 0 4 0 0.0 0.2 0.4 0.6 0.8 1.0 1.2 nb/la ba/la continental and island arc basalts kangâmiut dykes clear lake basalts continental flood basalts fig. 10. ba/la versus nb/la showing the differences between arc basalts, the kangâmiut dykes, and basalts from the clear lake volcanic field. data for the arc field are from bailey et al. (1989), tormey et al. (1991), francalanci et al. (1993), pearce et al. (1995), bacon et al. 1997), gust et al. (1997), and kelemen et al. (2003). data for the continental flood basalt field are from hooper & hawkesworth (1993), lightfoot et al. (1993), wooden et al. (1993), peate & hawkesworth (1996), and storey et al. (1997). data for the clear lake volcanic field are from charles lesher (unpublished data). 0.8 1.0 1.2 1.4 1.6 1.8 2.0 2.2 1.0 1.5 2.0 2.5 3.0 3.5 la/smn dy/ybn 3.0 2.9 2.8 2.7 2.6 2.5 2.42.22.01.81.61.4 2% 4% 6% degre e of m eltin g s ol id us p re ss ur e (g p a) segregation pressure (gpa) fig. 11. chondrite-normalised dyyb versus la/sm ratios for the kangâmiut dykes and mantle melting models using the algorithm of fram & lesher (1993) based on a 0.5% depleted pm source composition, where f is the melt proportion. the model assumes partial melting proceeds by incremental non-modal batch melting at 1% per kbar of decompression in a corner flow melting regime. melts are pooled after each kbar of decompression. ree distribution coefficients are taken from green (1994). the garnet-spinel transition is modelled as a gradual change between 30 and 25 kbar. the spinel–plagioclase transition is modelled as a gradual change between 14 and 10 kbar. residues are recalculated after each melting increment and adjusted for pressure dependent phase transitions using melting reactions as given by fram & lesher (1993). model curves, for melting starting at 3.0, 2.9, 2.8, 2.7, 2.6 and 2.5 gpa and ending at 0.5 gpa, define a melting grid, where solid subvertical lines contour constant mean melt fraction, whereas the dashed subvertical lines contour final pooled melt segregation pressure (i.e. the top of the melting column). slightly modified from mayborn & lesher (2004). 75 occur during the flux of fluids from the slab into the mantle wedge (miller et al. 1994; pearce et al. 1995; becker et al. 1999). mafic rock suites in volcanic arc settings also typically contain a large proportion of high-alumina basalts (perfit et al. 1980; brophy & marsh 1986; kelemen et al. 2003). if the kangâmiut dykes are arc-related, they should show lile enrichment, hfse depletions, and high al 2 o 3 . figure 10 shows a comparison of the nb/la and ba/la ratios for the kangâmiut dykes, island arcs, continental arcs, and continental flood basalts. relative to arc basalts, the kangâmiut dykes have lower ba/la and nb/la ratios – unlike arc-related basalts. additionally, subduction zone basalts typically have al 2 o 3 contents of 19–15 wt% (plank & langmuir 1988, 1992; kelemen et al. 2003), whereas all of the kangâmiut dykes have lower al 2 o 3 concentrations (16–12 wt%). thus, the kangâmiut dykes have none of the geochemical characteristics of subduction related basalts, contrary to previous conjecture (cadman et al. 2001). a more detailed analysis of the cadman et al. (2001) hypothesis also raises significant questions about its viability. cadman et al. (2001) propose that the kangâmiut dykes formed after ridge subduction resulting in a ‘slab window’ passing beneath metasomatised mantle. the resulting mantle upwelling lead to melting within hydrated mantle wedge material. although cadman et al.’s hypothesis would explain elevated water contents postulated for kangâmiut dyke magmas, such an origin would also be expected to impart an arc geochemical signature to the magmas. it is instructive to directly compare the composition of the kangâmiut dykes with those from the clear lake volcanic field located in the coastal region of northern california and associated with the development of a slab window after passage of the mendocino triple junction (furlong & schwartz 2004). figure 10 compares the ba/la and nb/la ratios for clear lake basalts, with typical arc basalts, the kangâmiut dykes and continental flood basalts. it is evident from these, among other, geochemical indices that the kangâmiut dykes lack the expected arc signature postulated by cadman et al. (2001). rather the kangâmiut dykes have compositions consistent with their derivation from asthenospheric mantle supplying normal continental flood basalts. active rifting, plume and passive rifting hypotheses the temperature of the mantle is an important difference between plume associated rifting and passive rifting. mayborn & lesher (2004) presented a detailed analysis of the temperature of the mantle during kangâmiut dyke magma genesis as constrained by ree systematics. they used the algorithm of fram & lesher (1993), as shown in fig. 11, to propose that the kangâmiut dykes are the results of mantle melting with a mean solidus pressure of c. 2.75 gpa and a mean extent of melting of 5%. when compared to the solidus for nominally anhydrous mantle (fig. 12), this mean solidus temperature would correspond with a potential mantle temperature of 1420°c. this temperature estimate falls at the lower end of potential temperatures estimated for c. 2.0 ga mantle by richter (1984, 1420–1600°c) and abbott et al. (1994, 1380–1680°c) based on secular cooling models and geochemical data for precambrian morb-type basalts, respectively. addition1200 1400 1600 1800 0 1.0 2.0 3.0 4.0 5.0 6.0 7.0 pressure (gpa) temperature (°c) modern adiabat (tp = 1300°c) 2.0 ga adiabat (tp = 1420°c) 2.0 ga plume adiabats mantle so lidus fig. 12. plot showing the fertile peridotite solidus and adiabats for modern mantle, 2.0 ga mantle, and plume mantle associated with potential temperatures of 100 to 300°c greater than 2.0 ga mantle. pressures and temperatures for the fertile peridotite from hirschmann (2000). slightly modified from mayborn & lesher (2004). fig. 13. variation of k 2 o (in wt%) with zr (in ppm) for the kangâmiut dykes. solid lines show results of fractional crystallisation models using mineral proportions of 10:50:40 ol:pl:cpx and 40:40:20 hbl:pl:cpx. the model starting composition is the sample with lowest k 2 o and zr, which are 0.19 wt% k 2 o and 47 ppm zr. based on cadman et al. (2001). 0.0 0.4 0.8 1.2 1.6 0 100 200 300 400 zr k2o 10:50:40 ol:pl:cpx 40:40:20 hbl:pl:cpx 76 ally, an ambient mantle potential temperature of 1420°c at 2.0 ga. is consistent with constraints from continental freeboard that suggests that mantle temperatures were below 1430°c by the mid-archaean (galer 1991). thus, the 1420°c mantle temperature for the kangâmiut dykes is consistent with ambient mantle temperatures for that time, reducing the need for anomalously high potential temperatures commonly associated with plume magmatism. the explanation that best fits both the geochemical and field data is that the kangâmiut dykes formed by decompression melting in a rift environment under ambient mantle conditions. this conclusion implies that the dykes are the products of rifting of kenorland supercontinent between 2.1 and 2.0 ga (williams et al. 1991). the implications of these findings for the temperature of the palaeoproterozoic mantle, the occurrence of palaeoproterozoic mantle plumes, and for palaeoproterozoic continental crustal growth are discussed in mayborn & lesher (2004). fractionation of the kangâmiut dykes as noted previously, the range in mg-number (0.60–0.21) of the kangâmiut dykes shows that they are not in equilibrium with mantle peridotite and do not represent direct mantle melts. thus, even the most primitive sampled kangâmiut dyke represents an evolved magma. first order observations of the whole-rock and mineral data show that the dykes evolved by ol:cpx:pl or hbl:cpx:pl fractionation with the late-stage introduction of fe-ti oxides into the fractionating assemblage. windley (1970) and bridgwater et al. (1995) proposed that hornblende was a primary crystallising phase from the kangâmiut dyke magmas partly based on the occurrence of large hornblende crystals in chilled margins. however, our petrographical studies of the chilled margins show that these amphiboles grew in situ during the final stage of solidification (see fig. 3). this does not preclude the possibility that hornblende was a stable and fractionating phase at depth and thus influenced the composition of evolved kangâmiut dyke magmas prior to their emplacement. the main difference between hbl:cpx:pl and ol:cpx:pl crystallisation to explain the magmatic evolution of the dykes is the relative cotectic proportions of hornblende and olivine. in the projection from qtz (fig. 6a) the dyke compositions lie within both the hbl:cpx:pl and ol:cpx:pl phase volumes. if hornblende is a fractionating phase its cotectic proportion would be c. 0.40 (based on the projections from qtz, cpx, and plag), with plagioclase and clinopyroxene at 0.40 and 0.20, respectively. if olivine, and not hornblende, is a fractionating phase, the cotectic proportions would be c. 0.10 olivine, 0.50 plagioclase, and 0.40 clinopyroxene. these different cotectic proportions can be used to determine if hornblende or olivine was a fractionating phase by examining the partitioning behaviour of potassium. experimentally determined amphibole-basaltic melt kds for potassium range between 1 and 2 (green 1994). in contrast, the olivine-basaltic melt kd is c. 0.0005 (green 1994) for potassium between olivine and basaltic liquid. figure 13 shows the results of fractional crystallisation modelling for k 2 o and zr. the cotectic assemblage 40:40:20 hbl:cpx:pl gives a bulk distribution coefficient (d) of 0.52 for k and 0.094 for zr. the olivine-bearing assemblage, 10:50:40 ol:cpx:pl, gives a bulk d of 0.16 for k and 0.026 for zr. as shown in fig. 13, the hornblende-bearing assemblage underestimates the concentration of k 2 o and is not consistent with the trend defined by the dyke data, whereas the olivine-bearing assemblage provides a better fit to the data. thus, these relationships show that the kangâmiut dykes evolved by the fractionation of olivine, clinopyroxene, plagioclase and late stage fe-ti oxides, and that hornblende was not a significant fractionating phase at any stage of their evolution. support for this conclusion comes from the 0.8 gpa melting experiments of mayborn (2000) showing that the cotectic assemblage for a water-bearing kangâmiut dyke starting material is olivine, clinopyroxene, and plagioclase, but no hornblende. 0 100 200 300 400 500 zr ce 0 20 40 60 80 100 0.1 0.2 0.3 0.4 0.5 f = 1.0 fig. 14. variation of ce with zr (in ppm) showing the results of a fractional crystallisation model where f is the melt proportion. tic marks on model curve are drawn at 0.1 intervals of f. 77 origin of water in the kangâmiut dykes based on the presence of hornblende in the kangâmiut dykes, bridgwater et al. (1995) speculated that the parental magmas for the dykes were derived from a hydrous protolith during thrusting of an amphibolite facies terrain beneath the granulite facies terrain in the southern nagssugtoqidian orogen. although this model does offer an explanation for the proposed high water contents of the dykes, the presence of hornblende itself, often as reaction rims on clinopyroxene or poikilitic grains enclosing plagioclase and clinopyroxene, only argues for elevated water contents during final stages of crystallisation of the dykes. it is, therefore, possible that the high water contents sufficient to stabilise hornblende resulted solely from its enrichment during crystal fractionation. the modal abundance of hornblende in dykes not affected by nagssugtoqidian deformation is c. 5–20% for dolerites and 10–35% for dioritic centres. since amphiboles contain c. 2 wt% water, these modes would indicate a whole-rock water concentration of 0.1 wt% in the primitive dolerites and 0.7 wt% in the more evolved dioritic centres. whether these differences in concentrations between the primitive and evolved samples are related to enrichment during crystal-liquid fractionation can be evaluated using the following equation for fractional crystallisation: c l = c 0 × f (d–1) (1) where c 0 is the initial concentration, c l is the liquid concentration, f is the proportion of liquid, and d is the bulk distribution coefficient. danyuschevsky et al. (2000) showed that in mafic systems water will have a bulk distribution coefficient of c. 0.01. starting with a magma with 0.1 wt% (c 0 ) h 2 o, representing the primitive dolerites, and ending with 0.7 wt% (c l ) h 2 o, representing the evolved dioritic centres, requires 86% crystallisation (f = 0.14) of the primitive magma. figure 14 shows the relative enrichment of highly incompatible elements zr and ce during fractional crystallisation, where ce is used as a proxy for water given their similar incompatibilities (danyuschevsky et al. 2000). the amount of fractionation required to relate the primitive samples to dioritic centres is 0.3–0.13, corresponding to 70–87% crystallisation. thus, the evolved dykes reflect sufficient fractionation to explain the difference in water concentrations between the primitive dykes and dioritic centres. the origin of the water in the primitive dykes can also be addressed using equation 1. the most primitive kangâmiut dykes (mgo > 6 wt%) have a modal hornblende content of 5–10% indicating a maximum of 0.1–0.2 wt% water content in the rocks. this suggests that the amount of water in the parental magma derived from the mantle is less than 0.2 wt%. equation 1 can also approximate fractional mantle melting and can help constrain the amount of water in the mantle source needed to produce a primitive magma containing 0.2 wt% h 2 o. in this case, the unknown variable is c 0 , the initial concentration in the mantle. the concentration in the liquid (c l ) is 0.2 wt%, and d is equal to 0.01. the evaluation of ree systematics, presented by mayborn & lesher (2004) and illustrated in fig. 11, shows that the average f value for mantle melting leading to the kangâmiut dykes was 0.05. thus, using these values in equation 1 results in a concentration in the mantle source (c 0 ) of 0.01 wt% (100 ppm). this is well within the range of 28–300 ppm h 2 o given by bell & rossman (1992) for the upper mantle containing nominally anhydrous phases. as such, the water present in the hornblende within the kangâmiut dykes can be reasonably accounted for given estimates of its original concentration in primary melts and enrichment through subsequent differentiation. although these considerations do not rule out bridgwater et al.’s (1995) model for the kangâmiut dykes derived from an amphibolite facies protolith, we show that differentiation of partial melts derived from depleted upper mantle can readily explain the occurrence of late crystallising hornblende in the evolved kangâmiut magmas. the kangâmiut dykes and the nagssugtoqidian orogeny the preservation of both igneous and metamorphic features in the kangâmiut dyke swarm provides an excellent opportunity to evaluate the amount of crustal thickening that likely occurred during the nagssugtoqidian orogeny. determining the amount of thickening requires knowledge of the depths associated with emplacement and peak metamorphism for currently exposed dykes. field relationships show brittle deformation of host rocks and segmentation of the kangâmiut dykes into en échelon arrays during emplacement. reches & fink (1988) proposed that the segmentation of dykes into en échelon arrays occurs when they cross from the ductile into the brittle regime. in modern continental crust the brittle– ductile transition is observed as the seismic to aseismic transition at depths of 10–15 km (chen & molnar 1983). chen & molnar (1983) and williams (1996) give temperature estimates for the brittle–ductile transition between 450 and 250°c. 78 fahrig & bridgwater (1976) presented palaeomagnetic data from dykes and host rocks unaffected by nagssugtoqidian metamorphism, and showed that the host rocks and dykes record different declinations. these differences in declination show that the host rocks were below their curie temperature during dyke emplacement. fahrig & bridgwater (1976) do not discuss the magnetic carrier in the host rocks, but an examination of host rock samples suggests that the magnetic carrier(s) are magnetite and/or ilmenite. the curie temperatures of these minerals vary due to solid solutions amongst magnetite-ulvöspinel and hematite-ilmenite, but the upper limit is 580°c if the magnetic carrier is pure magnetite. additional support for dyke emplacement into host rocks with temperatures below 580°c comes from 40ar/ 39ar dating of dykes and host rocks. willigers et al. (1999) presented 40ar/39ar cooling ages from dykes in the southern foreland that gave a mean age of 2.02 ga. this age is within error of the 2.04 ga emplacement age determined by dating of igneous zircons (nutman et al. 1999). the 40ar/39ar cooling age of a regional granitic host rock is 2.5 ga (willigers et al. 1999). this older age indicates that the host rocks have remained below 480°c, the closure temperature of argon in hornblende (harrison 1981), since 2.5 ga. field evidence of brittle deformation, palaeomagnetic data, and 40ar/39ar cooling ages all indicates that the peak crustal temperatures of exposed basement hosting the kangâmiut dykes were less than c. 450°c at the time of dyke emplacement. estimates of the geothermal gradient appropriate for continental crust at 2.0 ga can help to constrain the depth of the 450°c isotherm and thus the depth of dyke emplacement. the geotherm is computed from the heat flow equation assuming an exponential distribution of heat producing elements that includes contributions from heat conduction, advection, and production (carslaw & jaeger 1959): where t is temperature in °c, q* is the reduced heat flow at the crust-mantle boundary in mwm–2, z is the depth in km, k is the thermal conductivity of the crust in wm–1k– 1, ao is the concentration of heat producing elements at the earth’s surface in µwm–3, d is crustal thickness in km, and hr is the length scale for the decrease in heat producing elements with depth in km. current average values are q = 30 mwm–2, k = 2.25 wm–1k–1, ao = 0.75 µwm– 3, d = 35 km, and hr = 15 km. the model geotherm shown in fig. 15 uses these values with the exception of ao = 1.2 to account for higher heat production during the palaeoproterozoic and archaean (stein 1995). using the geotherm in fig. 15 and a maximum host rock tempera(2)0 0.3 0.6 0 200 400 600 800 10 20 d ep th ( km ) t (°c) 0 geotherm p (g pa ) fig. 15. model continental geotherm constructed using the heat flow equation shown in text (equation 2). solid vertical line represents the 450°c isotherm that intersects the geotherm at c. 0.3 gpa (10 km) as shown by dashed horizontal line. 0 0.2 0.4 0.6 0.8 1.0 1.2 200 400 600 800 ky and sil kangâmiut dyke peak metamorphism p (gpa) t (°c) g ra nu lit e a m ph ib ol ite fig. 16. pressure versus temperature diagram showing the intersection of the amphibolite/granulite facies transition with the kyanite– sillimanite transition. the black rectangle around the intersection shows possible temperature and pressure ranges of peak metamorphism of the kangâmiut dykes and metasediments in ikertooq fjord. the al 2 sio 5 phase diagram is from holdaway (1971) and the amphibolite to granulite transition is based on the first occurrence of orthopyroxene in experiments on mafic rocks by spear (1981). t = q* z k + a o d 2 k 1 − e −z / hr( ) 79 ture of 450°c require that the dykes intruded to a minimum depth of 10 km corresponding to a lithostatic pressure of c. 0.3 gpa. this estimate of a shallow emplacement level for the dykes is also supported by the clustering of whole-rock data near the low-pressure ol:cpx:pl cotectics as previously shown in fig. 7. application of the clinopyroxene geothermobarometry developed by putirka et al. (1996) provides another independent estimate of the depth of dyke emplacement. this geothermobarometer uses compositions of clinopyroxenes and their host rocks to estimate the pressure and temperature of clinopyroxene crystallisation. there are two important assumptions when applying this geothermobarometer to chilled margins in the kangâmiut dykes. first, we assume that the whole-rock composition is a close approximation of the original liquid composition. second, we assume that the cores of the clinopyroxene phenocrysts were once in equilibrium with this liquid. one test of equilibrium between the whole-rock and clinopyroxene phenocrysts is given by the feo/mgo ratios in the whole rocks and pyroxenes. the feo/mgo ratios in the clinopyroxenes from eight chilled margins and the feo/mgo ratio in their host rocks yield an average k d fe-mg for clinopyroxene and liquid of 0.32 ± 4. based on experimental work, k d fe-mg for basaltic systems between 1 atm and 1.5 gpa ranges from 0.22–0.36 (baker & eggler 1987; putirka et al. 1996). our estimates for the kangâmiut dykes fall within this range. applying the putirka et al. (1996) geothermobarometer to clinopyroxene phenocrysts cores and whole-rock compositions from chilled margins of eight kangâmiut dykes gives temperatures of 1199–1170°c and pressures of 0.75– 0.35 gpa. the upper pressure limit of 0.75 gpa indicates a maximum recorded depth of fractionation recorded by clinopyroxene phenocrysts of c. 25 km. the lower pressure estimate of 0.35 gpa constrains a maximum emplacement depth of c. 12 km, since the clinopyroxene phenocrysts in the chilled margins must have formed at a depth greater or equal to the final depth of dyke emplacement. this is consistent with the preceding results from temperature estimates of the host rocks during emplacement that indicate a maximum of c. 0.3 gpa or a depth of c. 10 km. if these independent estimates of the kangâmiut dyke emplacement depths are taken as representative of the dyke swarm in general, then a consideration of peak metamorphic conditions during the nagssugtoqidian orogeny can be used to constrain the amount of crustal thickening during orogenesis. the majority of the metamorphism of the northern portion of the swarm during the nagssugtoqidian orogeny occurred at amphibolite facies, with the exception of the northernmost portion within ikertooq fjord where the transition to granulite facies metamorphism occurs. the amphibolite to granulite facies transition is marked by the first appearance of orthopyroxene in mafic rocks and is known to occur at c. 800°c (spear 1981). the constraint on the pressure of the granulite facies metamorphism comes from the presence of kyanite-sillimanite paragneisses that are interleaved with sheets of dyke-bearing orthogneisses. the presence of granulite facies metamorphosed kangâmiut dykes and the alumina-silicate-bearing gneisses indicates that peak metamorphism occurred at conditions corresponding to both the amphibolite to granulite and kyanite to sillimanite transitions. figure 16 shows that these transitions indicate a peak metamorphic pressure of c. 0.9 gpa. this pressure is consistent with the results of mengel et al. (1995) who determined metamorphic pressures on the kangâmiut dykes in the ikertooq region using tweequ geothermobarometry (berman 1991). knowing the approximate depth of emplacement and the pressure of peak metamorphism provides constraints on the amount and style of burial during the nagssugtoqidian orogeny. emplacement at 0.3 gpa of pressure followed by peak metamorphism at 0.9 gpa requires an increase of 0.6 gpa. this indicates a minimum of 20 km of crustal thickening between dyke emplacement and peak metamorphism. a probable mechanism of crustal thickening in this case is thrust imbrication and crustal loading of material from north to south. the structural feature associated with this imbrication is most likely the ikertôq shear zone (fig. 1). the imbrication of rock types, the lithostratigraphic changes, including the disappearance of the kangâmiut dykes, and the lateral continuity of the ikertôq shear zone suggest that it is a major structure capable of accommodating displacement of material that buried the northern portion of the dyke swarm with 20 km of overburden. crustal thickening must have occurred over a minimum map distance of 50 km extending from the ikertôq shear zone to at least the itivdleq shear zone (fig. 1) and farther to the south approaching the nagssugtoqidian front where the last significant metamorphism occurs. summary and conclusions the 2.04 ga kangâmiut dyke swarm in west greenland is composed of tholeiitic dykes that intruded during passive rifting of archaean continental crust. the current level of exposure corresponds to emplacement depths less than 10 km based on estimated host rock temperatures less than 450°c during emplacement and geothermobarometry for 80 kangâmiut dyke clinopyroxenes. major and trace element systematics show that the parental magmas for the kangâmiut dykes differentiated by fractionation of plagioclase, clinopyroxene, olivine, and late state fe-ti oxides. the rare-earth element systematics of the dykes indicate initiation of mantle melting at c. 2.75 gpa, corresponding to a potential mantle temperature of c. 1420°c. this temperature is consistent with ambient mantle temperature estimates for 2.0 ga and shows that the kangâmiut dyke swarm formed during passive rifting of the kenorland supercontinent. anomalously hot plume mantle is not required for their generation. subsequent metamorphism of the northern portion of the swarm reached granulite facies, with an estimated temperature of 800°c and pressure of 0.9 gpa. the emplacement pressure of less than 0.3 gpa and peak metamorphism at 0.9 gpa indicate a minimum of 20 km of crustal thickening associated with the nagssugtoqidian orogeny. crustal thickening likely occurred during thrusting of material from the central nagssugtoqidian orogen southward over the southern nagssugtoqidian orogen along the ikertôq shear zone. acknowledgements we are especially grateful to the late david bridgwater, whose boundless energy and enthusiasm for the kangâmiut dykes inspired this work from beginning to end. we also thank flemming mengel, jim connelly, and minik rosing for their support of this project at various stages, and andy saunders and karen hanghøj for their constructive reviews of the final manuscript. this work was partially supported by the danish lithosphere centre and grants from the us national science foundation (ear 97-06677 and oce 98-11453). references abbott, d., burgess, l., longhi, j. & smith, w.h.f. 1994: an empirical thermal history of the earth’s upper mantle. journal of geophysical research 99(7), 13835–13850. arculus, r.j. 1994: aspects of magma genesis in arcs. lithos 33(1–3), 189–208. aspler, l.b. & chiarenzelli, j.r. 1998: two neoarchean supercontinents? 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(eds): mechanism of igneous intrusion, 79–92. liverpool: seel house press. wooden, j.l., czamanske, g.k., fedorenko, v.a., arndt, n.t., chauvel, c., bouse, r.m., king, b.w., knight, r.j. & siems, d.f. 1993: isotopic and trace-element constraints on mantle and crustal contributions to siberian continental flood basalts, noril’sk area, siberia. geochimica et cosmochimica acta 57(15), 3677–3704. 84 appendix dyke location latitude longitude trend thickness (m) samples 1 søndre strømfjord 66°05.266'n 053°33.211'w 010 2 430901 middle of dyke, 430902 near contact 2 søndre strømfjord 66°05.26'n 053°33.5'w 080 20 430903 ~7 m from dyke contact 3 søndre strømfjord 66°05.26'n 053°33.5'w 010 0.4 none 4 søndre strømfjord 66°05.26'n 053°33.7'w 020 ? none 5 itilleq 66°33.3'n 053°02.5'w 086 6 430904 ~2 m from contact 6 itilleq 66°33.3'n 053°02.5'w ? 1.5 430905 middle of dyke 7 ikertooq 66°58.1'n 052°28.9'w 079 0.3 none 8 ikertooq 66°58.1'n 052°28.9'w 123 0.15 none 9 ikertooq 66°57.55'n 052°31.7'w 010 2 430910 0.5 m from dyke margin 10 ikertooq 66°58.95'n 052°26.8'w 080 2 430915 middle of dyke 11 ikertooq 66°49.7'n 052°16.7'w 054 0.3 430917 middle of dyke 12 ikertooq 66°50.1'n 052°19.2'w ? ? 430919 middle of dyke 13 itilleq 66°32.25'n 052°45.0'w 080 8 430923 middle of dyke 14 itilleq 66°32.85'n 052°47.15'w ? 0.1 430926 whole width of dyke 15 itilleq 66°33.3'n 052°51.3'w 095 1 430929 middle of dyke 16 itilleq 66°33.03'n 052°53.54'w 035 8 430930 ~2.5 m from contact, 430931 dyke contact 17 itilleq 66°33.5'n 052°56.0'w 080 2 430933 0.3 m from contact 18 itilleq 66°34.7'n 052°56.0'w 062 10 430935 middle of dyke 19 itilleq 66°34.7'n 052°56.0'w 090 14 430936 middle of dyke 20 itilleq 66°35.1'n 052°49.0'w 090 8 430939 middle of dyke 21 itilleq 66°35.1'n 052°47.5'w 079 8 430940 middle of dyke 22 itilleq 66°34.75'n 052°48.5'w ? ? 430942 middle of dyke 23 itilleq 66°34.75'n 052°48.5'w ? ? 430943 middle of dyke 24 itilleq 66°33.1'n 053°04.0'w ? 25 430946 middle of dyke 25 itilleq 66°33.7'n 052°55.1'w 087 16 430948 middle of dyke, 430950 dyke contact 26 itilleq 66°33.25'n 052°38.5'w 100 1 430951 middle of dyke 27 itilleq 66°33.25'n 052°33.0'w 096 20 430952 middle of dyke 28 itilleq 66°32.6'n 052°27.0'w 057 30 430953 middle of dyke 29 itilleq 66°31.8'n 052°41.0'w 061 5 430955 ~1.5 m from contact, 430956 ~5 m from contact 30 itilleq 66°31.9'n 052°38.5'w 085 15 430957 dyke contact, 430959 ~5 m from contact 31 itilleq 66°31.8'n 052°36.2'w 065 18 430960 middle of dyke 32 itilleq 66°31.7'n 052°34.0'w 078 22 430961 middle of dyke 33 mouth of itilleq 66°29.9'n 053°33.5'w 022 20 430965 middle of dyke 34 mouth of itilleq 66°30'n 053°34.8'w 065 25 430966 ~7 m from contact 35 mouth of itilleq 66°30.05'n 053°35.8'w 117 40 430967 ~7 m from contact 36 mouth of itilleq 66°30.5'n 053°36.4'w 006 10 430969 middle of dyke 37 south of maniitsoq 65°22.7'n 052°47.3'w 345 4.5 430970 middle of dyke 38 east of maniitsoq 65°25.5'n 052°24.0'w ? 15 430972 middle of dyke 39 east of maniitsoq 65°25.4'n 052°23.0'w 054 0.4 430973 middle of dyke 40 east of maniitsoq 65°25.5'n 052°18.0'w 117 12 430974 middle of dyke 41 east of maniitsoq 65°26.0'n 052°14.8'w 354 7 430975 middle of dyke 42 east of maniitsoq 65°35.15'n 052°46.0'w 002 49 430977 dyke contact, 430979 1 m from contact, 430981 5 m from contact, 430982 17 m from contact 43 north of maniitsoq 65°39.85'n 052°37.0'w 003 25 430983 middle of dyke, 430984 dyke contact 44 north of maniitsoq 65°44.4'n 052°38.5'w 344 0.5 430986 middle of dyke 45 north of maniitsoq 65°38.75'n 052°37.5'w 002 16 430987 middle of dyke, 430988 ~6 cm from contact 46 north of maniitsoq 65°36.9'n 052°43.5'w 005 10 430990 ~4 m from contact, 430991 ~0.5 m from contact table a1. field data for kangâmiut dykes examined for this study 85 dyke location latitude longitude trend thickness (m) samples 47 north of maniitsoq 65°40.1'n 052°49.0'w 010 17 430992 middle of dyke 48 north-east of kangaamiut 65°54.9'n 053°14.5'w 050 12 430993 ~2 m from contact 49 north-east of kangaamiut 65°53.95'n 053°16.0'w 045 27 430994 middle of dyke 50 north-east of kangaamiut 65°53.4'n 053°14.95'w 033 1 none 51 north-east of kangaamiut 65°53.4'n 053°14.95'w 110 0.2 none 52 north-east of kangaamiut 65°53.4'n 053°14.95'w 030 1.5 430995 middle of dyke 53 north-east of kangaamiut 65°53.4'n 053°14.95'w 110 2 430996 middle of dyke 54 north-east of kangaamiut 65°52.6'n 053°14.0'w 063 60 430997 andesitic portion of dyke, 430998 mafic portion of dyke 55 north-east of kangaamiut 65°50.8'n 053°13.0'w 064 1.5 432101 middle of dyke 56 north-east of kangaamiut 65°50.8'n 053°13.0'w 015 15 430999 middle of dyke 57 søndre strømfjord 66°01.4'n 053°28.65'w 028 60 432103 middle of dyke 58 kangerluarsussuaq 66°17.5'n 053°05.65'w 065 40 432104 dyke contact, 432105 middle of dyke 59 kangerluarsussuaq 66°17.5'n 053°05.8'w 057 40 432107 middle of dyke 60 kangerluarsussuaq 66°17.25'n 053°07.5'w 022 1 432108 middle of dyke 61 kangerluarsussuaq 66°17.25'n 053°07.5'w 145 45 432106 1 m from contact, 432108 middle of dyke 62 kangerluarsussuaq 66°17.15'n 053°08.5'w 045 21 432111 middle of dyke 63 kangerluarsussuaq 66°17.05'n 053°11.25'w 028 40 432112 ~7 m from contact 64 kangerluarsussuaq 66°16.9'n 053°09.7'w 031 140 432115 dyke contact, 432116 ~2.5 m from w contact, 432118 ~45 m from w contact, 432119 ~70 m from w contact, 432120 ~95 m from w contact, 432121 ~19 m from e contact, 432122 ~18 m from w contact, 432136 ~40 m from w contact, 432137 middle of dyke 65 kangerluarsussuaq 66°39.5'n 053°03.0'w 065 50 432123 middle of dyke 66 east of itilleq 66°29.3'n 052°25.0'w 065 15 432125 middle of dyke 67 east of itilleq 66°29.1'n 052°24.0'w 050 8 432128 middle of dyke 68 east of itilleq 66°33.1'n 052°07.5'w 072 1 432129 middle of dyke 69 east of itilleq 66°31.75'n 052°18.0'w 050 4 432130 middle of dyke 70 east of itilleq 66°30.5'n 052°27.5'w 090 25 432138 middle of dyke 71 east of itilleq 66°27.9'n 052°27.0'w 022 70 432133 middle of dyke 72 east of itilleq 66°26.0'n 052°40.5'w 080 5 432134 middle of dyke 73 east of itilleq 66°27.45'n 052°45.5'w 055 4 432135 middle of dyke 74 kangerluarsussuaq 66°16.6'n 053°17.0'w 032 17 432139 middle of dyke 75 kangerluarsussuaq 66°15.3'n 053°22.5'w 036 18 432140 middle of dyke, 432151 9 m from n contact, 432152 6 m from n contact, 432153 4 m from n contact, 432154 2 m from n contact, 432155 1 m from n contact, 432156 0.5 m from n contact, 432157 0.25 m from n contact, 432158 dyke contact, highly jointed 76 kangerluarsussuaq 66°14.6'n 053°33.0'w 020 25 432143 middle of dyke, 432144 5 m from e contact, 432145 e contact 77 kangerluarsussuaq 66°14.7'n 053°31.7'w 044 0.4 432147 nw contact, 432148 middle of dyke 432149 se contact 78 kangerluarsussuaq 66°15.15'n 053°29.0'w 086 2 432150 middle of dyke 79 mouth of itilleq 66°30.05'n 053°35.8'w 090 50 430206 and 430207 middle of dyke 80 mouth of itilleq 66°30.05'n 053°35.8'w 090 15 430208 and 430209 middle of dyke 81 mouth of itilleq 66°30.05'n 053°35.8'w 090 ? 430210 middle of dyke 82 mouth of itilleq 66°30.05'n 053°35.8'w 090 0.1 430211 whole width of dyke 83 søndre strømfjord 66°01.4'n 053°28.65'w ? ? 430258 dyke contact 84 north-east of kangaamiut 65°52.6'n 053°14.9'w 022 30 430267 dyke contact, 430265 middle of dyke 85 north-east of kangaamiut 65°53.6'n 053°14.9'w 045 30 430283 dyke contact, 430284 middle of dyke 86 north-west of kangaamiut 65°56.6'n 053°28.0'w 25 60 158074 and 430288 dyke contact, 432102 and 158077 middle of dyke table a1 (continued) 86 maniitsoq 25 km hamborgerland kangaamiut sø nd re st røm fjo rd sukkertoppen iskappe itilleq itilleq sisimiut ikertooq ta sersuaq sarfartoq maligiaq avalleq evigh eds fjor d søn dre iso rto q 65°30' 66° 66°30' 51° 53° 52° 1370135 6 12 34 7 8 9 10 11 12 1417 24 19 16 15 20 26 21 18 22 28 23 27 25 30 3129 35 323436 37 33 38 39 40 41 56 46 45 44 43 42 47 4849 51 54 55 50 5352 1357 13581359 136013611362 1364 1363 1365 1366 1367 1368 1369 1371 13721373 1374137513761377 1378 79 80 81 82 86 85 84 83 fig. a1. map showing the locations of dykes examined in this study. 71 the arctic region is warming more rapidly than the global average (amap 2017) and it is well established that this warming is at least partially responsible for the greenland ice sheet losing mass at an accelerating rate, raising concern worldwide (e.g. kahn et al. 2015; rahmstorf et al. 2015). it is essential to monitor the changes of the greenland ice sheet to be able to assess the potential environmental, social and economic implications around the globe, and to provide decision-makers with reliable data. the annual mass-budget deficit of the greenland ice sheet has grown over the past two decades due to increases in surface melting (van den broeke et al. 2017) and ice-flow acceleration (kahn et al. 2015). currently, and for the last two decades, the greenland ice sheet is the single largest arctic cryospheric contributor to global sea-level rise and the greenland ice-surface melt rates are projected to increase as the arctic continues to warm (amap 2017). the snowline is here defined as the maximum elevation during the melt season at which snow remains from the previous accumulation season (cogley et al. 2011). the snowline is a valuable climate indicator as its position integrates the competing effects of melt (increasing snowline elevation) and snow accumulation (decreasing snowline elevation). thus the snowline provides a key holistic variable indicating climate change. we have developed a methodology that determines snowline elevation utilising the moderate resolution imaging spectroradiometer (modis) sensor on the terra satellite. the modis sensor produces a global dataset on a daily basis, with a resolution varying between 250 m and 1 km, in 36 bands covering the visible to thermal wavelengths. using modis, we derived the maximum snowline altitude for the greenland ice sheet for the years 2000–2017. we are producing a freely available, consistent dataset that provides an important tool for the monitoring of the long-term impact of climate change on the greenland ice sheet. direct comparison with field observations from automatic weather stations (awss) from the programme for monitoring of the greenthe greenland ice sheet – snowline elevations at the end of the melt seasons from 2000 to 2017 robert s. fausto and the promice team* * signe b. andersen, andreas p. ahlstrøm, dirk van as, jason e. box, daniel binder, michele citterio, william colgan, konstanze haubner, karina hansen, nanna b. karlsson, kenneth d. mankoff, allan ø. pedersen, anne solgaard and baptiste vandecrux 500 km thu upe kpc tas kan nuk sco qas egp land 185.3 370.6 555.9 741.2 926.5 1112 1297 1482 1668 1853 2038 2224 2409 2594 2779 2965 3150 snowline 2016 promice gimp dem fig. 1. greenland map showing the location of promice automatic weather stations and the 2016 snowline as derived from terra satellite data using the moderate resolution imaging spectroradiometer (modis) sensor. the locations of the ground-control automatic weather stations (promice) are indicated. egp: east grip. kan: kangerlussuaq. kpc. kronprins christian land. nuk: nuuk. qas: qassimiut. sco: scoresby sund. tas: tasiilaq. thu: thule. upe: upernavik. dem: digital elevation model. gimp: greenland ice mapping project. © 2018 geus. geological survey of denmark and greenland bulletin 41, 71–74. open access: www.geus.dk/bulletin http://www.geus.dk/bulletin 7272 land ice sheet (promice) network validates the snowline dataset derived from modis. we use the services of the cryoclim internet portal, providing an operational and permanent service for long-term systematic climate monitoring of the cryosphere, to distribute our snowline product. more specifically, end-of-melt season, 1 km2 resolution raster grids illustrating snow and bare-ice surfaces, and snowline shape files can be downloaded via cryoclim. here, we describe the snowline classification algorithm, its validation and its interannual variations for 18 years spanning 2000–2017. snowline classification algorithm we processed all modis mod12km and mod03 scenes covering greenland from late july to the beginning of september 2000–2017. we used the surface-type detection algorithm of fausto et al. (2015) that distinguishes between bareice and snow surfaces. fausto et al. (2015) uses normalised thresholds (th) from calibrated radiances (mod021km) between the near-infrared band 5 (1230–1250 nm) and the visible band 10 (483–493 nm) with surface-type thresholds thdry snow≤0.86, 0.86<thmelting snow<0.94 and thglacier ice≥0.94. the classification algorithm is updated by implementing a new surface-type threshold to address a ‘noisy’ snow classification over the northern part of the ice sheet identified by fausto et al. (2015). the algorithm is thus supplemented by thbare ice>265–2.1×lat, where lat is latitude. thbare ice is defined as: t h bare ice= c 0 + c1×b1+ c 2 ×b2+ c 3 ×b3 + c 5 ×b5 + c7×b7 where c0=–0.0015, c1=0.160, c2=0.291, c3=0.243, c5= 0.112, c7=0.081 and b1 to b 7 designate band 1 to band 7. cloud-covered regions are removed using the mod35_l2 dataset. subsequently pixels are classified for every modis scene as either snow or bare ice for the whole greenland ice sheet. daily classification scenes are aggregated to yield a maximum extent of bare ice to define an end-of-melt-season snowline. snowlines from peripheral glaciers are generally excluded, and the snowline products are based on an algorithm success rate of over 95% classified pixels. validation to help validate the modis data we make use of the promice automatic weather station network that currently consists of two or three stations primarily in the ablation area in eight ice sheet regions. each automatic weather station records a suite of meteorological and glaciological measurements, supplemented by e.g. surface-height changes due to accumulation or ablation (fig. 1; van as et al. 2016). to validate the classified snowline elevation at the end of the melt season, we use the mass-budget values from the promice weather stations (fig. 2; e.g. fausto et al. 2012) at different elevations to calculate the vertical surface mass-balance gradient for all eight promice transects to determine the equilibrium line altitude (ela, zero mass budget), for direct comparison with modis estimated snowline elevation (fig. 1). aws balance profiles from the upernavik region, and those indicating an ela above 2000 m are excluded as we find them unrealistic. the location of the upper aws should be close to the actual ela to get the best balance profiles. in total, we exclude 25% or 17 out of 67 balance profiles. fig. 2. south greenland promice automatic weather station at the end of the 2013 melt season. from the stakes to the left of the weather station, the amount of melt (c. 4 m) is directly visible. the melt is also measured with a pressure transducer system drilled into the ice. 73 figure 3 illustrates the performance of the modis end-ofmelt-season snowline algorithm for all promice regions in greenland. the correlation (r=90%, p=0.0001, n=50) and the root-mean-square error (rmse=200 m) are reasonable as the ela and snowline elevation can be different due to superimposed ice formation (cogley 2011). the mean difference between snowline altitude and ela is −104 m. results and discussion figure 1 illustrates the location of remotely sensed snowline plotted on top of the digital elevation model (dem) from the greenland ice mapping project (gimp, howat et al. 2014). the snowline is easily visible in the southern, western, and northern parts of greenland due to the relatively even terrain, while the snowline shows a more complicated pattern in the mountainous terrain in east greenland (fig. 1). the snowline separates bare ice from snow areas and can therefore be used to document the change in bare-ice areas. we find the extent of bare-ice exposure to be increasing in the period 2000–2017 at an average rate of c. 500 km2 per year (fig. 4), which roughly corresponds to the size of the danish island of bornholm. this increase in the bare-ice area is insignificant, but it demonstrates a small average gain of melt over accumulation since 2000. the increasing trend in the bare-ice area is consistent with increasing greenland mass loss due to surface processes (van den broeke et al. 2017). both independent, in situ observations (machguth et al. 2016) and remotely sensed observations (hall et al. 2012; tedesco et al. 2017) show that the greenland melt area is expanding to higher elevations. further, the increase in bare ice enhances the positive feedback mechanism of a darkening ice sheet surface (ice is darker than snow), which affects the surface mass and energy balance of the greenland ice sheet (box et al. 2012). figure 4 also illustrates the inter-annual variability of the 2000–2017 snowlines, which is highly dependent on the complicated seasonal weather systems around greenland. for instance, the below average snowline of the snowy year of 2016/2017 is consistent with positive albedo anomalies that reduced melting in 2017 (tedesco et al. 2017). uncertainties associated with the different surface-type detection are assessed with the elas derived from the aws surface mass-budget observations. figure 3 shows a significant correlation between the modis snowline and elas derived independently from promice awss. a reason for the difference between the two can be that the modis data have a spatial resolution of 1 km2, pan-ice sheet coverage and quasi-daily temporal coverage, while the footprints of the in situ measurements are small (5–50 m2), and surface patchiness is clear in aerial photography (stroeve et al. 2006). fausto et al. (2015) discuss an august anomaly in their monthly surface-type data set during the 2010–2014 period, illustrated by a noisy melting-snow classification in the northern ice sheet, which was most likely due to false classification. however, with the updated bare-ice threshold, we improve the detection of snow and ice surfaces (fig. 1), visualised by a less noisy snow classification of snow in the northern part of the ice sheet, resulting in a more reliable climate indicator for greenland. conclusions remotely sensed modis data can yield daily, automated classification of the greenland ice sheet surface type (snow and ice). validation indicates a high correlation (0.9) between modis-derived snowline altitudes and elas estimated from in situ measurements. the end-of-melt-season 1000 1500 1700 1500 1300 1100 900 700 500 automatic weather station ela (m a.s.l.) kpc sco tas qas nuk kan upe thusn ow lin e el ev at io n (m a .s. l.) 500 2004 2009 2014 190 000 170 000 150 000 130 000 110 000 90 000 70 000 year ar ea (k m ) 1999 2 bare-ice area y = 504x – 879029 fig. 3. the end-of-melt-season snowline elevation for 2000 to 2017 from modis vs. the promice aws-derived equiblibrium line altitude (ela). the blue line gives the 1:1 relation. the locations of the promice automatic weather stations are shown in fig. 1. fig. 4. end-of-melt-season bare-ice area for the greenland ice sheet for the years 2000–2017. 7474 snowline is useful as an ice-sheet climate indicator for the competing processes of surface accumulation and ablation, quantified by an average annual increase of c. 500 km2 of the bare-ice area for the 2000–2017 period. acknowledgements the programme for monitoring of the greenland ice sheet (promice) is funded by the geological survey of denmark and greenland (geus) and the danish ministry of energy, utilities and climate under the danish cooperation for environment in the arctic (dancea), and is conducted in collaboration with dtu space, denmark’s national space institute, and asiaq (greenland survey). the nuk and kan stations were/are (co-)funded by the greenland climate research centre (gcrc) and the greenland analogue project (gap), respectively. this study was funded by dk esa-prodex under the cryoclim project. references amap 2017: snow, water, ice and permafrost in the arctic (swipa) 2017, 269 pp. oslo: arctic monitoring and assessment programme (amap). box, j.e., fettweis, x., stroeve, j.c., tedesco, m., hall, d.k. & steffen, k. 2012: greenland ice sheet albedo feedback: thermodynamics and atmospheric drivers. the cryosphere 6, 821–839. cogley, j.g., hock, r., rasmussen, l.a., arendt, a.a., bauder, a., braithwaite, r.j., jansson, p., kaser, g., moller, m., nicholson, l. & zemp, m. 2011: glossary of glacier mass balance and related terms. ihp-vii technical documents in hydrology 86, iacs contribution 2. paris: unesco-ihp. fausto, r.s., van as, d. & promice project team 2012: ablation observations for 2008–2011 from the programme for monitoring of the greenland ice sheet (promice). geological survey of denmark and greenland bulletin 26, 25–28. fausto, r.s. et al. 2015: greenland ice sheet melt area from modis (2000–2014). geological survey of denmark and greenland bulletin 33, 97–100. hall, d.k., comiso, j.c., digirolamo, n.e., shuman, c.a., key, j.r. & koenig, l.s. 2012: a satellite-derived climate-quality data record of the clear-sky surface temperature of the greenland ice sheet. journal of climate 25, 4785–4798. howat, i.m., negrete, a. & smith, b.e. 2014: the greenland ice mapping project (gimp) land classification and surface elevation datasets. the cryosphere 8, 1509–1518. http://dx.doi.org/10.5194/tc-8-15092014 khan, s.a., aschwanden, a., bjørk, a.a., wahr, j., kjeldsen, k.k. & kjær, k.h. 2015: greenland ice sheet mass balance: a review. reports on progress in physics 78, 046801, 26 pp. http://dx.doi.org/10.1088/00344885/78/4/046801 machguth, h., macferrin, m., van as, d., box, j.e., charalampidis, c., colgan, w., fausto, r.s., meijer, h.a.j., mosley-thompson, e. & van de wal, r.s.w. 2016: greenland meltwater storage in firn limited by near-surface ice formation. nature climate change 6, 390–393. http:// dx.doi.org/10.1038/nclimate2899 rahmstorf, s., box, j., feulner, g., mann, m., robinson, a., rutherford, s. & schaffernicht, e. 2015: exceptional twentieth-century slowdown in atlantic ocean overturning circulation. nature climate change 5, 475–480. http://dx.doi.org/10.1038/nclimate2554 stroeve, j.c., box, j.e. & haran, t. 2006: evaluation of the modis (mod10a1) daily snow albedo product over the greenland ice sheet. remote sensing of environment 105, 155–171. http://dx.doi. org/10.1016/j.rse.2006.06.009 tedesco, m., box, j.e., cappelen, j., fausto, r.s., fettweis, x., hansen, k., mote, t., sasgen, i., smeets, c.j.p.p., van as, d., velicogna, i. & van de wal, r.s.w. 2017: greenland ice sheet: contribution to the arctic report card 2017. van as, d., fausto, r.s., cappelen, j., van de wal, r.s.w., braithwaite, r.j., machguth, h. & promice project team 2016: placing greenland ice sheet ablation measurements in a multi-decadal context. geological survey of denmark and greenland bulletin 35, 71–74. van den broeke, m.r., box, j.e., fettweis, x., hanna, e., noël, b., tedesco, m., van as, d., van de berg, w.j. & van kampenhout, l. 2017: greenland ice sheet surface mass loss: recent developments in observation and modelling. current climate change reports 3, 345–356. http://dx.doi.org/10.1007/s40641-017-0084-8 authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: rsf@geus.dk. http://dx.doi.org/10.5194/tc-8-1509-2014 http://dx.doi.org/10.5194/tc-8-1509-2014 http://dx.doi.org/10.1088/0034-4885/78/4/046801 http://dx.doi.org/10.1088/0034-4885/78/4/046801 http://dx.doi.org/10.1038/nclimate2899 http://dx.doi.org/10.1038/nclimate2899 http://dx.doi.org/10.1038/nclimate2554 http://dx.doi.org/10.1016/j.rse.2006.06.009 http://dx.doi.org/10.1016/j.rse.2006.06.009 http://dx.doi.org/10.1007/s40641-017-0084-8 mailto:rsf@geus.dk geological survey of denmark and greenland bulletin 35, 2016, 39-42 39 a complete log-stratigraphical breakdown of the middle ordovician to lower silurian shale-dominated succession is presented for the bornholm–skåne–kattegat area in southernmost scandinavia. a wireline log zonation developed for the onshore bornholm palaeozoic shales is extended to include the offshore palaeozoic shales in the adjacent rønne graben. a complete log zonation scheme for the cyrtograptus shale (late llandovery–wenlock) and the lower part of the colonus shale (ludlow) is presented including correlation within the bornholm–skåne–kattegat area. the cyrtograptus shale in the bornholm area is estimated to be 400 m thick and marks the shift to a rapidly subsiding foreland basin, heralding the caledonian orogeny. the lower palaeozoic shales in denmark and southernmost sweden are locally very thick (>3 km). the great thickness is mainly due to the presence of expanded silurian units that formed in an active foreland basin related to the caledonian orogeny to the south and west (michelsen & nielsen 1991; vejbæk et al. 1994, eriksson 2012; calner et al. 2013). the thick silurian succession has only been known from a few deep exploration wells of older date with very limited wireline logging information (cf. vejbæk et al. 1994). however, recent interest in the scandinavian lower palaeozoic shales as a potential shale gas resource has led to a number of new drill holes with modern geological and geophysical logs, considerably increasing our knowledge. in a recent summary, schovsbo et al. (2015) presented a review of the scientific work carried out over the last decade on bornholm, based on the complete drilling and coring of the onshore lower palaeozoic succession. in may 2015, data acquired as part of the exploration programme for shale gas in skåne by royal dutch shell plc. (shell) were released by the swedish authorities, adding to our knowledge on the lower palaeozoic shales. these newly released data are integrated here with log data from wells onand offshore bornholm as well as older deep wells in western denmark. the oderup c4-1 well the exploration programme carried out by shell in skåne included drilling of the lövestad a3-1, hedeberga b2-1 and oderup c4-1 wells (fig. 1) and acquisition of some 80 km of 2d seismic profiles (calner & pool 2011; pool et al. 2012). the oderup c4-1 well drilled in central skåne has the most complete wireline log suite and is used here as a key well for establishing a silurian well log stratigraphy (fig. 2). for all three wells eriksson (2012) presented a detailed stratigraphic and lithologic evaluation as well as a wireline log-correlation between skåne and bornholm. below a brief description is presented of the drilled sequence in the oderup c4-1 well, mostly based on calner & pool (2011) and eriksson (2012). lower palaeozoic strata were encountered below a quaternary cover between 34.5 and 926.9 m terminal depth and the well terminated in the lower cambrian hardeberga formation. thin lower cambrian læså (0.3 m) and gislöv formations (0.6 m) are overlain disconformably by the middle cambrian – lower ordovician (tremadocian) alum shale formation, which is 76.3 m thick middle–upper ordovician and silurian stratigraphy and basin development in southernmost scandinavia niels h. schovsbo, arne t. nielsen and mikael erlström 50 km bornholm r ø n n e g ra b en skåne sweden germany kattegat c4-1 terne-1norwegian–danish basinringkøbing–fyn high denmark a sommerodde-1 pernille-1 a3-1 b2-1 lower palaeozoic strata caledonian front well with lower palaeozoic strata fig. 1. distribution of lower palaeozoic strata in southernmost scandinavia with location of wells north of the caledonian front that reach the lower palaeozoic. only wells referred to in the paper are named. wells drilled by shell in skåne are lövestad a3-1, hedeberga b2-1 and oderup c4-1. © 2016 geus. geological survey of denmark and greenland bulletin 35, 39–42. open access: www.geus.dk/publications/bull 4040 (836–912.3 m). the alum shale formation is in turn disconformably overlain by the lower middle ordovician komstad limestone formation (832.5–836 m). the komstad limestone is in turn overlain by 44 m of middle and upper ordovician shales (788.5–832.5 m) representing the almelund, sularp, mossen, fjäcka and lindegård formations (fig. 2). the sularp, mossen and fjäcka formations are here collectively referred to as the dicellograptus shale. the base of the rastrites shale (uppermost ordovician – lower silurian) is marked by a change to dark lithologies at 788.5 m, which is also recorded as high gamma ray (gr) readings (fig. 2). in the oderup c4-1 final well report, the top of the rastrites shale was identified at 733.5 m, where a change from dark to carbonaceous light grey shale occurs. however, adopting the definitions of schovsbo et al. (2015) the top of the rastrites shale is here defined at a slightly deeper level corresponding to the inflection point of the formation resistivity log at 749 m (fig. 2). the top of the cyrtograptus shale is here defined at 460 m where a change from the dark grey cyrtograptus shale to the light grey colonus shale occurs. the same horizon has a distinct motif on the gr log being marked by low readings (fig. 2). the colonus shale is dominated by light grey to green-grey, micaceous and slightly calcareous to arenaceous shale with frequent intercalations of grey limestone nodules (calner et al. 2013). wireline log-correlations the log-stratigraphy of palaeozoic shales on bornholm, originally defined by pedersen & klitten (1990), has recently been redefined and extended to include the complete rastrites shale (log zones f1–f5) and the onshore part of the cyrtograptus shale (log zones g1–g5) by schovsbo et al. (2015). the same log zones have been identified in the oderup c4-1, terne-1 and pernille-1 wells (fig. 3). correlations between the terne-1 well and sections on bornholm and between the oderup c4-1 well and bornholm have previously been presented by michelsen & nielsen (1991) and eriksson (2012), respectively, and only minor updates and corrections are made here. in the correlation panels (fig. 3) additional log zones have been defined in the cyrtograptus and colonus shales (new log zones g6–g7 and h1–h2, respectively; see below). the new log zones of the cyrtograptus and colonus shales the silurian g5 log zone (see schovsbo et al. 2015) was not completely penetrated by the sommerodde-1 well onshore bornholm and the top of the zone is here defined in the pernille-1, oderup c4-1 and terne-1 wells as a characteristic gr peak co-occurring with a change to slightly higher sonic log velocities (fig. 3). according to bjerreskov (1993) and vejbæk et al. (1994), the base of the pernille-1 well is within the lower wenlock (uppermost sheinwoodian), which is slightly older than the cyrtograptus shale with intercalated tuff-bearing sandstone (homerian; upper wenlock) that is exposed at the sommerodde beach locality onshore bornholm (bjerreskov & jørgensen 1983). this beach section immediately overlies the succession 75125 0 250 0 500 1/vp, ms/ft 0 100 200 300 400 500 600 700 800 900 c o lo n u s sh al e gr, api r o rd . fu m c l c u p p er s ilu ri an l o w er s ilu ri an c yr to gr ap tu s sh al e a lu m s h gr, api m ea su re d d ep th b el o w g ro u n d s u rf ac e (m ) cba cba q l r, ohm.m 10 1000 scale change (casing no logs) 3 2 4 1 fig. 2. a: gamma ray curve. b: interval transit time velocity. c: formation resistivity in the oderup c4-1 well, skåne. note that the definition of the rastrites shale follows schovsbo et al. (2015). lc: lower cambrian. mc: middle cambrian. fu: furongian. ord: ordovician. q: quaternary. alum: alum shale fm. l: lindegård fm. r: rastrites shale. 1: hardeberga, læså and gislövs fms. 2: komstad limestone fm. 3: almelund fm. 4: dicellograptus shale (sularp, mossen and fjäcka fms). 41 penetrated by the nearby sommerodde-1 well. it is inferred that the tuff-bearing sandstone interval exposed at sommerodde is seen in the log pattern in the pernille-1 well as zones with low gr response and high sonic velocities reflecting cemented sandstone beds (fig. 3). the g6 log zone is characterised by generally upward-decreasing gr values and increasing sonic velocities. the top of the zone is defined in the wells by a slight drop in gr activity. in the oderup c4-1 well the top occurs just below a distinct low gr and high sonic velocity bed (fig. 3). the gr log continues to show a gradual decrease throughout log zone g7. the top of the zone is defined at an increase in gr values and a general change to more fluctuating log signatures as seen in the oderup c4-1 well. the h1 log zone is distinguished by an increase in gr values in comparison with the g7 zone. the top is placed at a narrow and significant gr peak with low sonic velocities that could represent bentonite beds. the h2 log zone sill casing f5 f4 e2 e1 f3 g7 g6 g5 g4 g1 d1 d2 d3 f1 f2 3300 3200 3100 3000 2900 2800 2700 2600 2500 2400 2300 oderup c4-1 sommerodde-1pernille-1 / r as tr it es s h al e l . a lm . t ø ye n a lu m s h al e d ic . 3000 3100 3200 3300 3400 3500 3600 1/vp, ms/ft d ep th b el o w g ro u n d s u rf ac e (m ) 0 100 200 300 400 500 600 700 800 900 0 10000 1000 m ea su re d d ep th b e lo w r t ( m ) 0 100 200 f5 f4 e f3 h2 h1 g7 g6 g5 g4 g3 g2 g1 d r o tl . c o lo n u s sh al e gr, api terne-1 0 200 125 50 100 m f1/2 a d l c yr to gr ap tu s sh al e c yr to gr ap tu s sh al e r as tr it es c ar b o n ife ro u s z ec h st ei n gr, api gr, api gr, api gr, api gr, api 10000 0 200 1/vp, ms/ft 125 50 0 200 1/vp, ms/ft 125 50 m ea su re d d ep th , b el o w r t ( m ) lo fu mc fu lc mc scale change d ep th b el o w g ro u n d s u rf ac e (m ) d ep th b el o w g ro u n d s u rf ac e (m ) scale change scale change fig. 3. well log correlation panel between the terne-1, oderup c4-1, pernille-1 and sommerodde-1 wells (for location, see fig. 1). the lower log zonation (d1–d3, f1–f5, g1–g5) is adopted from schovsbo et al. (2015); the g5 (top)–g7 and h1–h2 log zones are new. for simplicity not all log zones are shown including those in the alum shale (cf. schovsbo et al. 2015). ras. sh: rastrites shale. l: lindegård fm. d: dicellograptus shale. alm: almelund shale. rotl: rotliegendes. rt: rotary table. 4242 is characterised by a variable log pattern that reflects the more variable lithologies in this unit. the transition from passive margin to foreland basin the transition from passive margin to a foreland basin occurred in the early silurian (vejbæk et al. 1994; eriksson 2012; calner et al. 2013). in the kattegat area continuous sedimentation took place throughout this interval whereas bornholm and skåne were characterised by temporary local uplifts and condensations (fig. 3). during most of the early, mid and early late ordovician, bornholm was uplifted and no sedimentation occurred apart from the bioclastic early middle ordovician komstad limestone fm, whereas skåne, as exemplified by the oderup c4-1 well, shows variable degrees of condensation and uplift during the late early ordovician to early silurian. from the onset of deposition of the cyrtograptus shale, the difference in sedimentation between the areas converges (fig. 3). this shift was also noted by eriksson (2012) and calner et al. (2013) and is interpreted to reflect the development of a rapidly subsiding foreland basin. the foregoing uplifts thus heralded the later subsidence and are seen as isostatic adjustments of the margins of baltica, perhaps including a foreland bulge passage. the uplift first affected bornholm and later skåne but apparently did not influence the kattegat area where continuous subsidence occurred throughout the ordovician and silurian. conclusions a complete log-stratigraphical breakdown of the middle ordovician to lower silurian shale-dominated succession is presented for the bornholm–skåne–kattegat area in southernmost scandinavia. the cyrtograptus shale (of late llandovery–wenlock age) is subdivided into seven log zones and is estimated to be approximately 400 m thick in the bornholm area based on the sommerodde-pernille composite section (fig. 3). the lower part of the overlying colonus shale (ludlow) is divided into two log zones; the total thickness of this shale is unknown. variable degrees of condensation and uplift occurred in skåne and on bornholm during the late early ordovician to early silurian. from the onset of the cyrtograptus shale deposition, the difference in sedimentation between the areas converges and this shift is interpreted to reflect the development of a rapidly subsiding foreland basin. acknowledgements geocenter denmark is thanked for financial support to the project ‘silurian stratigraphy and basin development’ awarded to nhs and atn. comments from reviewers svend stouge and mikael calner helped improving the final version of the manuscript. references bjerreskov, m. 1993: pernille-1 well graptolite fauna. geus report file 10567. reported as part of a efp-89 project. dgu kunderapport 1–5. bjerreskov, m. & jørgensen, k.a. 1983: late wenlock graptolite-bearing tuffaceous sandstone from bornholm, denmark. bulletin of the geological society of denmark 31, 129–149. calner, m. & pool, w. 2011: the first deep wells in the lower palaeozoic colonus shale trough, sorgenfrei-tornquist tectonic zone, southern sweden. gff 133 , 58–59. stockholm: geological society of sweden. calner, m., erlström, m., lehnert, o. & ahlberg, p. 2013: lower palaeozoic geology of southern sweden. in: calner, m. et al. (eds): the lower palaeozoic of southern sweden and the oslo region, norway. field guide for the 3rd annual meeting of the igcp project 591. sgu rapporter och meddelanden 133, 6–9. eriksson, m. 2012: stratigraphy, facies and depositional history of the colonus shale trough, skåne, southern sweden. dissertations in geology at lund university 310, 37 pp. michelsen, o. & nielsen, l.h. 1991: well records on the phanerozoic stratigraphy in the fennoscandian border zone, denmark: hans-1, sæby-1, and terne-1 wells. danmarks geologiske undersøgelse serie a 29, 37 pp. pedersen, g.k. & klitten, k. 1990: anvendelse af gamma-logs ved korrelation af marine skifre i vandforsyningsboringer på bornholm. dansk geologisk forening årsskrift 1987–89, 21–35. pool, w., geluk, m., abels, j. & tiley, g. 2012: assessment of an unusual european shale gas play: the cambro-ordovician alum shale, southern sweden: proceedings of the society of petroleum engineers/european association of geoscientists and engineers unconventional resources conference, vienna, austria, march 20–22, 152339. schovsbo, n.h., nielsen a.t. & klitten, k. 2015: the lower palaeozoic now fully cored and logged on bornholm. geological survey of denmark and greenland bulletin 33, 9–12. vejbæk, o.v., stouge, s. & poulsen, k.d. 1994: palaeozoic tectonic and sedimentary evolution and hydrocarbon prospectivity in the bornholm area. danmarks geologiske undersøgelser serie a 34, 21 pp. authors’ addresses n.h.s., geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: nsc@geus.dk a.t.n., dept. of geosciences and natural resource management, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. m.e., geological survey of sweden (sgu), kiliansgatan 10, se-223 50 lund, sweden. geological survey of denmark and greenland bulletin 35, 2016, 43-46 43 injection of chemically tuned, ‘smart’ water in oil reservoirs may increase both oil recovery rates and the total recovery (e.g. morrow & buckley 2011; austad 2013; zeinijahromi et al. 2015). this kind of water management has gained increased importance in the danish north sea reservoirs due to decreasing sweep efficiency in maturing oilfields. knowledge about the compatibility of the injected water with local formation waters is, however, a prerequisite for successful implementation. here, we present a regional overview of formation waters from oil reservoirs in the danish north sea, which comprise three main types of formation brine, and one type of modified seawater related to extensive water flooding. the water types show a distinct geographical distribution, which reflects original connate waters that are modified by saline brine being either depleted or enriched in so4 2–. formation water and produced water database in order to characterise the water types we have selected a total of 33 water analyses, 25 of produced water and six of formation water from north sea wells (three core samples from the francisca-1 well and production tests from the boje-1, elna-1 and m-9x, wells) and finally two analyses of seawater (north sea mean water composition and a treated low-sulphate seawater), see fig. 1. for characterisation, samples analysed for na, k, ca2, mg2, sr2, ba2, cl–, and so4 2– were used. water density had been measured for most of the samples, however, it was estimated for four samples. the data were collected from samuelsen et al. (2009), mackay et al. (2012), and undall-behrend (2012) and from final well reports for the boje-1, elna-1, francisca-1 and m-9x wells. water type classification to classify the water types in our database, principal component analysis (pca) was applied, whereby a matrix x of measured data (n samples, p variables) is transformed into sets of projection subspaces delineated by principal components (each a linear combination of all p variables), which display variance-maximised interrelationships between variables (esbensen 2010, esbensen et al. 2015). pca score plots display groupings, or clusters, of samples based on compositional similarities, as described by the variable correlations (shown in accompanying loading plots). they also quantify the proportion of total dataset variance that can be modelled by each component, see fig. 2. all data analyses in this work are based on auto-scaled data. the data analysis was performed in two steps. step one is a pca analysis of all 33 samples to investigate relationships between seawater and reservoir water (fig. 2a, b). based hereon, pure seawater and the samples produced from skjold, dan b, dan f and halfdan, which represent extensively seawater-flooded reservoirs, were removed types of formation water and produced water in danish oiland gasfields: i mplications for enhanced oil recovery by injection of ‘smart’ water niels h. schovsbo, hanne d. holmslykke, claus kjøller, kathrine hedegaard, lars kristensen erik thomsen and kim h. esbensen valdemar a kraka regnar dan fdan b cecilie lulita gorm roar rolf m-9 x svend skjold boje-1 siri nini dagmar halfdan tyra se s. arne tyra w tyra e valdemar b fransisca-1 harald e harald w valdemar a kraka regnar dan fdan b cecilie lulita gorm roar rolf m-9 x svend skjold boje-1 siri nini elna-1 dagmar halfdan tyra se s. arne tyra w tyra e valdemar b fransisca-1 harald e harald w uk germany norway denmark the netherlands 500 km 0 50 km well-head well field deliniation fig. 1. location of the wells used for water-type characterisation in the danish part of the north sea. © 2016 geus. geological survey of denmark and greenland bulletin 35, 43–46. open access: www.geus.dk/publications/bull 4444 to examine the relationships between the primary water types, which is done in step 2 of the analysis (fig. 2c, d). in both pca models the first two principal component axes resolve 80% of the total data variance, with the main trend expressed on the pca-1 axis being salinity variation (seen as high positive pca-1 loadings for cl, na and k). the pca-2 axis displays high positive loadings of so4 2– and high negative loadings of elements such as ba, sr, mg and ca.this could indicate that so4 2– concentrations control the concentration of ba2+, sr2+, mg2+ and ca2+ due to the low solubility of e.g. barite (baso4) and anhydrite (caso4). however, chemical speciation calculations using the numerical code phreeqc suggest sub-saturation of so4 2–-bearing minerals except for barite in all water types. composition and occurrence of water types from the pca analysis, four water types can be identified based on natural groupings in the pca-2 versus pca-1 plot (fig. 2). the most likely cause of the salinity variation is a variable mixing of primary connate waters with brine originating from permian zechstein salt, which may or may not be so4 2–-rich (warren et al. 1994). the characteristics and occurrence of each of the water types are presented below. water types 1–3 have compositional characteristics similar to types presented in the comprehensive overview paper by warren et al. (1994). water type 1 plots within a very narrow group in the third quadrant of the pca-2– pca-1 plot (fig. 2c), characterised by low salinity water with an overall low abundance of all elements (fig. 3). this water type is found in the boje1, francisca-1, roar, tyra e, tyra se and valdemar fields and thus occurs in a broad range of reservoirs from lower to upper cretaceous – paleocene chalk to oligocene sandstone. the location borders the greater tyra–valdemar area, in geographical areas separated from saline permian brines (fig. 4). the water is interpreted to reflect unmodified so4 2–-depleted formation water. fig. 2. principal component analysis, pca. a: score. b: loading plot of pca on all data. c: score. d: loading plot of pca from which seawater from the dan b, dan f, skjold and halfdan installations and produced waters were excluded. the plot models 79% (a, b) and 83% (c, d) of the total data variance, respectively; variance proportions are shown along each component axis. water types are classified according to their groupings in a (water type 4) and c (water type 1, 2, 3). f-1: francisca-1. vald: valdemar. sw: seawater. lssw: low so4 2–-treated seawater. p c a 2 ( 1 6 % ) -0.4 0.0 0.4 0.8 -4 0 4 8 pca 1 (63 %) 0.0 0.2 0.4 p c a 2 ( 1 6 % ) -2 0 2 4 p c a 2 ( 1 6 % ) -2 0 2 4 pca 1 (63 %) pca 1 (67%) -4 0 4 8 0.0 0.2 0.4 pca 1 (67%) -0.4 0.0 0.4 0.8 p c a 2 ( 1 6 % ) boje-1 f-1roar tyra e tyra se tyra w valda&b dagmar harald w kraka lulita m-9x regnar rolf elna-1 harald east s. arne siri a siri c siri d gorm na k ca mg sr ba cl so4 density siri b svend na ca mg sr ba cl so4 density lssw dan b dan f halfdan skjold kdagmar harald wkraka lulita m-9x regnar rolf elna-1 harald e s. arnesiri a siri c siri b siri d svend tyra se vald a vald b gorm f-1 b d chalk field sandstone field produced water formation water sea water c a 45 water type 2 is characterised by positive pca-1 and negative pca-2 scores (fig. 2c) and can compositionally be characterised by medium to high salinities, no so4 2–, and high to very high ca and ba concentrations (fig. 3). this water type occurs in the harald e and w, lulita, s. arne, siri, nini, stine and cecilie fields, all of which are located in the northern part of the danish north sea and in the siri canyon – i.e. in reservoirs that range in age from jurassic to paleocene and both in chalk and sand lithologies. water type 2 is interpreted to reflect formation water modified by so4 2– depletion. water type 3 plots with generally positive pca-1 and pca-2 scores in fig. 2c, reflecting medium to high salinities and variable, low to high so4 2– concentrations (fig. 3). this water type is found in the dagmar, elna-1, gorm, kraka, m-9x, regnar, rolf and svend fields, most clearly expressed in the intensely fractured dagmar field sample. this field is situated on top of a salt dome that has reservoir oil in chalk and zechstein carbonates. type 3 waters are restricted to chalk reservoirs overlying salt domes in the southern salt dome province, and are interpreted as formation water enriched in so4 2–. water type 4 plots close to, or together with seawater with negative pca-1 and positive pca-2 scores in fig. 2a, corresponding to low to medium salinity with high so4 2– concentrations (fig. 3). this water type occurs in the dan, halfdan and skjold fields and is interpreted to be the result of decades of extensive water flooding performed by the operator (energistyrelsen 2013). analyses of water from the dan field (the m-9x well; fig. 2c) prior to flooding suggest that it was originally filled with water type 3. implications for enhanced oil recovery by injection of ‘smart’ water water injection is currently applied in several of the danish oil fields, mainly in order to provide pressure support. however, such injection may also have secondary effects such as increased imbibition, alteration of the reservoir rock wettability or mobilisation of fines with a resulting increase in reservoir sweep. in some cases, the specific chemical composition of the injection water may be important. thus, it has been suggested that carbonate rocks become more water wet if the injection water contains so4 2– in combination with excess ca2+ or mg2+ (e.g. austad 2013). the result is enhanced oil recovery, which is even more pronounced both for chalk and sandstone if the salinity of the injection water is significantly lower than that of the formation water (morrow & buckley 2011; austad 2013). na k cl ca mg sr ba s e a w at e r n o rm al is e d , p p m /p p m 0.01 1 100 10 000 na k cl ca mg sr ba so4so4 water type 1 water type 2 water type 3 water type 4 a b fig. 3. a: water concentration normalised to seawater composition. b: calculated average compositions of the four water types observed in this study. for display purposes, ba2+ and sr2+ concentrations of 0 ppm in seawater have been increased to 1 ppm. water types: 1: low salinity 2: high salinity so4 3: high salinity + so4 4: sea water-modified type 2-dominated type 3-dominated type 1-dominated type 2-dominated type 3-dominated type 1-dominated type 2-dominated type 3-dominated type 1-dominated 0 50 km fig. 4. occurrence of resolved water types in the danish oiland gasfields. for location names see fig. 1. the water types are geographically restricted and reflect both structural basin development and reservoir conditions. the seawater modified water type 4 is assumed to have originated as water type 3 based on pre-waterflooding formation water analysis and its structural position within the salt dome province. 4646 although several different mechanistic explanations have been suggested, a supposed change in carbonate rock wettability would involve surface chemical reactions such as ion exchange between so4 2– and oil molecules (e.g. austad 2013). following this argumentation, it is likely that injection of ‘smart’ so4 2–-bearing water in chalk reservoirs would have the largest effect in reservoirs with saline formation water depleted in so4 2– (water type 2). however, the application of so4 2–-rich water in reservoirs with this type of connate water is not straightforward, as there is a risk of scaling and subsequent clogging of the reservoir if the injected water is mixed with the connate water, due to its high concentrations of ca2+, ba2+, and sr2+ (samuelson et al. 2009; mackay et al. 2012). another risk related to injection of so4 2–-bearing water in so4 2–-depleted reservoirs is the possibility of hydrogen sulphide formation due to so4 2– reducing microbial activity. for reservoirs already enriched in so4 2– (water type 3) or with water of relatively low salinity (water type 1), other types of injection water may have greater effects on oil recovery. in shaly sand reservoirs, injection of low-salinity ‘smart’ water can also mobilise clay fines, in order to intentionally clog current flow paths and redirect the flow in the reservoir (e.g. morrow & buckley 2011; zeinijahromi et al. 2015). in this case, the mobilisation of non-swelling clays is provoked solely by the change in salinity. therefore, application of this type of water technology seems to be most relevant in reservoirs with connate water of relatively high salinity, such as most of the reservoirs in the siri canyon (water type 2). conclusions four water types are present in the danish north sea: so4 2–-bearing, mediumto highly saline water (type 3), so 4 2–-depleted medium to high saline water (type 2), so4 2–-depleted low saline (type 1), and a seawater-modified manifestation (type 4 water). these water types reflect variable mixing of connate water with deeper brines and are tied in with the known hydrocarbon provinces. type 2 represents the siri canyon and the south arne – svend areas. water type 3 is characteristic of salt dome reservoirs, while water type 1 represents the greater tyra–valdemar area. the variable chemistry of the formation water in the danish north sea imposes regional differences in production strategies and hence in the designing of ‘smart’ water for enhanced oil recovery. the classification of water types presented here shows that their composition is predictable and related to geographical domains in the north sea. this may be useful when designing procedures for optimal water management in the danish north sea, e.g. application of low salinity water flooding on mature fields, or in some cases even during the exploration stage. references austad, t. 2013: water-based eor in carbonates and sandstones: new chemical understanding of the eor potential using ‘smartwater’. in: sheng, j.j. (ed.): enchanced oil recovery field case studies, 301–335. waltham, ma, usa: elsevier. energistyrelsen, 2013: danmarks olieog gasproduktion 2013, 105 pp. københavn: energistyrelsen. esbensen, k.h. 2010: multivariate data analysis – in practice. an introduction to multivariate data analysis and experimental design, 5th edition, 598 pp. oslo: camo software as. esbensen, k., schovsbo, n.h. & kristiansen, l. 2015: down-hole permeability prediction – a chemometric wire-line log feasibility study from a north sea chalk well. geological survey of denmark and greenland bulletin 33, 13–16. mackay, e., ginty, w.r. & jones, t.j. 2012: oilfield scale management in the siri asset – paradigm shift due to the use of mixed pwri / seawater injection. 74th eage conference and exhibition incorporating europec 2012. copenhagen, denmark, 4–7 june 2012. spe 154534, 1–12. morrow, n. & buckley, j. 2011: improved oil recovery by low-salinity waterflooding. journal of petroleum technology 63, 106–112. samuelsen, e.h., frederiksen, r.a., heath, s.m., thornton, a., sim, m., arefjord, a. & mcara, e.k. 2009: downhole scale control through continuous injection of scale inhibitor in the water injection – a field case. conference tekna geilo paper 240309, 23 pp. undall-behrend, g. 2012: produceret vand på tyra øst f. bachelorprojekt århus maskinmesterskole, 75 pp. warren, e.a., smalley, c.p. & howarth, r.j. 1994: compositional variations of north sea formation waters, part 4. geological society, london. memoirs 15, 119–208. zeinijahromi, a., ahmetgareev, v., badalyan, a., khisamov, r. & bedrikovetsky, p. 2015: case study of low salinity water injection in zichebashskoe field. journal of petroleum science research 4, 16–31. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark; e-mail: nsc@geus.dk geological survey of denmark and greenland bulletin 13, 2007, 21-24 modern geological research into the late and postglacial history of the inner danish waters (i.e. kattegat, bælthavet and øresund, plus the adjoining fjords and estuaries) began at the turn of the last century. since then most investigations have focused on the timing of the initial marine inundation of the area, the early to mid-holocene changes in land–sea confi guration and sea level changes during the mid-holocene littorina period. research on the late holocene marine environment has received less emphasis, undoubtedly due to problems in finding continuous marine sediment records, as sedimentation in large areas of the danish waters seems to have been characterised by complex spatial and temporal patterns of deposition and non-deposition (e.g. lykke-an dersen et al. 1993). in an ongoing project we aim to ex plore the continuous development of da nish coastal environments over the last 9000 years using a variety of proxy data, including molluscs, dia toms, foramini fera, algal pigments, plant macrofossils and physical properties of sedi ments. the project spans both environmental and cultural history, and addresses the important links between them, as the nature of the coastal environment has exerted major influences on cultural and societal expression and activity from meso lithic to mod ern times. this paper presents some of the first results from the project concerning environmental changes in the roman warm period (c. 2000–1600 years b.p.) as shown by changes in molluscan faunas at two coring sites in horsens fjord and tem pelkrog in southern isefjord (fig. 1). hydrography the present-day circulation pattern in the inner danish waters is dominated by a two-layer estuarine flow, driven by outflow of low-salinity surface water from the baltic sea and environmental change in danish marine waters during the roman warm period inferred from mollusc data peter rasmussen, kaj strand petersen and david b. ryves © geus, 2007. geological survey of denmark and greenland bulletin 13, 21–24. available at: www.geus.dk/publications/bull 21 fig. 1. map of denmark showing the location of the two study sites horsens fjord and tem pelkrog in isefjord and the present day sea surface salinities (psu; annual mean) in the danish waters. the red dots indicate fjords and estuaries with iron age shell middens. (modified from dahl et al. 2003). bælthavet includes storebælt and lillebælt and the sea between the islands south of fyn and sjælland. inflow of high-salinity bottom water from the north sea and skagerrak. due to high inflows of freshwater to the baltic sea from rivers there is a strong surface-salinity gradient from west to east: from >30 psu (practical salinity units, equivalent to ppt) in skagerrak decreasing eastwards to <5 psu in the baltic sea (fig. 1; al-hamdani et al. 2007 – this volume). there is also a strong salinity stratification (a halocline) within the water column, with a wedge of higher salinity, north sea-derived water underlying less dense, less saline surface water. material and methods two sediment core lengths of 6 m and 13.5 m were extracted from respectively horsens fjord and tempelkrog, both at a water depth of around 5 m and both consisting of homogeneous clay-gyttja. the two sediment records were accelerator mass spectrometry (ams) 14c-dated using terrestrial plant material, thus avoiding marine-derived material which suffers from uncertain reservoir and hard-water effects that can cause serious dating problems in danish fjord and marine waters (heier-nielsen et al. 1995). the horsens fjord and tempelkrog sediment cores were subsampled at intervals of 1 cm and 2 cm, respectively, which is equivalent to a resolution of 10–15 years in both records for the time period discussed in this paper. molluscs were extracted from the cores by wet-sieving of known sediment volumes through a sieve with a mesh size of 0.1 mm. macrofaunal specimens were identified and counted and species numbers were calculated for 100 ml of sediment (fig. 2). results and discussion molluscan faunas based on analogies with their present-day ecological requirements in relation to salinity, temperature, depth and substrate, quaternary molluscs are useful tools as indicators of environmental and climatic changes through time. changes in surface salinity (and temperature) have often been linked with fluctuations in different marine populations, for example stocks of fish and molluscs, which tolerate only certain ranges of salinity and temperature. in the transitional area between the skagerrak and the baltic sea, the salinity gradient seems to be the main limiting factor in the geographic distri bution of mollusc species (sorgenfrei 1958; petersen 2004). 22 fig. 2. summary stratigraphic, macrofaunal data from two sediment cores from horsens fjord and tempelkrog. abundances of mol luscs are expressed as specimens per 100 ml of fresh sediment. remains of echinoids are shown as present/not-present. the dotted lines delimit the roman warm period. note the change of x-axis scale. the summary stratigraphic, macrofauna data from hor sens fjord and tempelkrog during the period 2500 to 1500 years b.p. are shown in fig. 2. at both sites there is a distinct increase in mollusc abundance at the beginning of the roman warm period between c. 2000 and 1850 years b.p. and at tempelkrog again at the end of the roman warm period. at horsens, this increase is mainly due to mytilus edulis (blue mussel), rissoa albella, r. albella var. sarci, bittium reticulatum, cerastoderma edule (common cockle) and mysella bidentata, and at tempelkrog, hydrobia ulvae, h. ventrosa, h. neglecta and rissoa albella. at both sites, this increase in mollusc abundance is accompanied by the appearance of ostrea edulis (european flat oyster) and a more steady presence of remains of echinoids (sea urchins), both indicating more saline conditions than in the period before or after (fig. 2). ostrea edulis needs at least 25 psu in order to reproduce (jensen & spärck 1934) and is not present in the inner danish waters today, except for the western part of lim fjorden which is connected to the north sea. furthermore, the reproductive success of ostrea edulis is very sensitive to temperature with optimal conditions around 20–22°c (spärck 1924). echinoids are generally stenohaline and therefore disappear when the water becomes brackish (c. 10 psu). in addition, other mollusc species, which also indicate an increase in salinity to >25 psu, appear at horsens fjord, although with low abundance. these include parvicardium scabrum, abra nitida, rissoa violacea, gari fervensis, velutina velutina, odostomia umbilicaris and abra prismatica (sorgen frei 1958). the two study sites are located 120 km apart and the similar pattern in macrofaunal assemblages, with an increase in mollusc abundance and the appearance of a suite of almost fully marine taxa at various times during the roman warm period, thus testifies to a widespread change to more saline and productive conditions across the inner danish waters. the fairly regular presence of ostrea edulis (see below) may also suggest an increase in sea surface temperature during this period. the reason for two (or three) ‘episodes’ with changes in faunal compositon at tempelkrog and only one at horsens fjord is unclear but could be a result of spatial or temporal differ ences in substrate, local current conditions or postmortem (taphonomic) processes. sea and society inferences of higher salinity and productivity during the roman warm period from the palaeoenvironmental data are also supported by an independent line of evidence from the archaeological record. large shell middens dated to the centuries around 2000 years b.p., and predominantely composed of mytilus edulis, cerastoderma edule, littorina littorea (common periwinkle) and to a lesser extent ostrea edulis, are recorded along several danish and north german fjords (figs 1, 3; anger 1973; harck 1973; poulsen 1978; petersen 1985). the synchroneity between the inferred changes in the marine environment and the appearance of these iron age shell middens strongly suggest a causal connection between the two, implying that people responded to the increased productivity in the marine environment by a comprehensive and targeted gathering and processing of shellfish. outlook the evidence of salinity increase in the inner danish waters during the roman warm period only seems to be explicable by the more frequent inflow of high-salinity north sea water, which travelled through the danish straits and further east into the baltic sea. this scenario is supported by the presence of ostrea edulis in iron age shell middens as far into the baltic as the head of flensburg fjord and eckernförder bucht 23 fig. 3. section of an iron age shell midden at horsens fjord. the iron age middens are mainly composed of shells, charcoal and pot boilers (stones used in cooking) and are interpreted as specialised coastal sites used for gathering and processing of shellfish (poulsen 1978). mollusc analysis from north german middens suggests that they are seasonal sites used in the late summer or autumn (anger 1973). photograph courtesy of karen løkkegaard poulsen. (north germany; anger 1973; harck 1973); this is only possible through a combination of a more frequent input of high-salinity water and a higher rate of water exchange than today. the increased inflow of north sea water seems to have penetrated as far as the central baltic sea where diatom and isotopic data also suggest a salinity rise during the roman warm period (emeis et al. 2003). diatom-based sea-surface salinity reconstructions from southern skagerrak also indicate higher salinity during the period (hebbeln et al. 2006), in good agreement with our results from the danish waters and evidence from the gotland basin (emeis et al. 2003). future work work is ongoing to analyse the other proxy records from the two sites over the last 9000 years to provide additional independent environmental information and to test the inferences made from the subfossil mollusc data. for instance, a model to relate fossil diatom assemblage composition to past surface-water salinity is being developed using a large range of contemporary samples from the western baltic sea, lim fjorden and coastal brackish lakes and fjords (the molten project: http://craticula.ncl.ac.uk/molten/jsp/; ryves et al. 2004). coastal sediments contain important natural archives of past environmental changes and palaeoecological techniques can provide a powerful means of revealing the natural variability of the marine environment and the links between environmental and socio-cultural changes over time. further more, this approach can establish the nature of environmental conditions in nearshore marine areas prior to the impact of modern society on coastal regions. it is especially important to establish ecological baseline conditions before realistic goals for environmental management of coasts can be set. acknowledgements the ongoing research project denmark’s coastal environment over the last 9000 years: linking cultural and hydrographic change is co-financed by the danish research council for the humanities (fkk) and the danish natural science research council (fnu), whose support is gratefully acknowledged. additional funding for the project has been provided by loughborough university, uk. karen løkkegaard poulsen is thanked for information about iron age shell middens. references al-hamdani, z.k., reker, j., leth, j.o., reijonen, a., kotilainen, a.t. & dinesen, g.e. 2007: development of marine landscape maps for the baltic sea and the kattegat using geophysical and hydrographical para meters. geological survey of denmark and greenland bulletin 13, 61–64. anger, k. 1973: untersuchungen an eisenzeitlichen muschelhaufen an der flensburger förde. offa 30, 55–59. dahl, k., lundsteen, s. & helmig, s.a. 2003: stenrev – havets oaser, 104 pp. copenhagen: gads forlag. emeis, k.-c., struck, u., blanz, t., kohly, a. & voss, m. 2003: salinity changes in the baltic sea (nw europe) over the last 10 000 years. the holocene 13, 411–421. harck, o. 1973: eisenzeitliche muschelhaufen an der schleswigschen ostund westküste. offa 30, 40–54. hebbeln, d., knudsen, k.-l., gyllencreutz, r., kristensen, p., klitgaardkristensen, d., backman, j., scheurle, c., jiang, h., gil, i., smelror, m., jones, p.d. & sejrup, h.-p. 2006: late holocene coastal hydrographic and climate changes in the eastern north sea. the holocene 16, 987–1001. heier-nielsen, s., heinemeier, j., nielsen, h.l. & rud, n. 1995: recent reservoir ages for danish fjords and marine waters. radiocarbon 37, 875–882. jensen, a.s. & spärck, r. 1934: bløddyr ii. saltvandsmuslinger. dan marks fauna 40, 208 pp. lykke-andersen, h., knudsen, k.l. & christiansen, c. 1993: the quaternary of the kattegat area, scandinavia: a review. boreas 22, 269–281. petersen, k.s. 1985: det sydfynske arkipelag. dets geologiske udvikling med særlig hensyntagen til havniveauændringer og den marine molluskfauna. in: skaarup, j. (ed.): yngre stenalder på øerne syd for fyn, 15–27. rudkøbing: langelands museum. petersen, k.s. 2004: late quaternary environmental changes recorded in the danish marine molluscan faunas. geological survey of denmark and greenland bulletin 3, 268 pp. poulsen, k.l. 1978: eisenzeitliche muschelhaufen in dänemark. offa 35, 64–85. ryves, d.b., clarke, a.l., appleby, p.g., amsinck, s.l., jeppesen, e., landkildehus, f. & anderson, n.j. 2004: reconstructing the salinity and environment of the limfjord and vejlerne nature reserve, den mark, using a diatom model for brackish lakes and fjords. canadian journal of fisheries and aquatic sciences 61, 1988–2006. sorgenfrei, t. 1958: molluscan assemblages from the marine middle miocene of south jutland and their environments. danmarks geo logiske undersøgelse ii. række 29, 356–503. spärck, r. 1924: undersøgelser over østersens (ostrea edulis) biologi i limfjorden, særlig med henblik paa temperaturens indflydelse paa kønsskiftet, 82 pp. københavn: centraltrykkeriet. authors’ addresses p.r. & k.s.p., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: per@geus.dk d.b.r., department of geography, loughborough university, loughborough, leicestershire, le11 3tu, uk. 24 geological survey of denmark and greenland bulletin 20, 2010, 67–70 67 the fiskenæsset complex in southern west greenland is part of the north atlantic craton and is a layered intrusion consisting of gabbro, ultramafic and anorthositic rocks that was deformed during multiple episodes of folding and metamorphism (myers 1985). we collected late-stage magmatic hornblenditic dykes and adjacent anorthosites and studied these samples integratively with several in situ techniques to determine the igneous and metamorphic history of the fiskenæsset complex. the work presented here is part of an ongoing joint project between the greenland bureau of minerals and petroleum and the geological survey of denmark and greenland (geus). here we report on new radiometric ages and mineral chemistry data from anorthosites from the north atlantic craton in southern west greenland (fig. 1). geological setting of majaqqap qaava despite the intense archaean deformation of the fiskenæsset area the original stratigraphy of the fiskenæsset complex was established from detailed field work at majaqqap qaava (myers 1985). towards the top part of the fiskenæsset complex hornblenditic pegmatite pipes cut the generally anorthositic and leucogabbroic layering. at majaqqap qaava, subvertical hornblenditic dykes are interpreted as representing a late magmatic stage of activity within the fiskenæsset complex (myers 1985). the fiskenæsset complex is surrounded and intruded by younger, mainly tonalitic gneisses (2.87–2.85 ga; næraa & scherstén 2008) that typically occur as felsic sheets intruded parallel to the magmatic layering. regional amphibolite-facies metamorphism affected the rocks at majaqqap qaava, but no granulite-facies metamorphism was recorded for this part of the fiskenæsset complex (myers 1985). zircon dating for u/pb zircon age determination we selected sample ggu 508216, which consists of hornblenditic dyke material and the surrounding anorthosite rock. the sample was crushed, sieved, and washed on a wilfley table. the zircons were hand-picked from the heavy mineral fraction, mounted in epoxy resin and polished. age determination was carried out by laser ablation inductively coupled mass spectrometry using an element2 and newwave 213 nm uv-laser system at geus following the procedures described in frei & gerdes (2009). the results are shown in fig. 2. the zircon spot analyses yielded a wide age span, with 207pb/206pb ages ranging from 2.70 ± 0.03 to 2.95 ± 0.03 ga (fig. 2a; 50 concordant grains out of 54). the oldest zircon grains in our sample are c. 2.95 ga (fig. 2a), which probably represents the intrusion age of the anorthosite complex. new data by a. polat and co-workers are consistent with this interpretation, as they obtained a sm-nd isochron age of c. 2.97 ga and a 207pb/206pb age of 2.95 ga for the intrusion of zircon record of the igneous and metamorphic history of the fiskenæsset anorthosite complex in southern west greenland nynke keulen, tomas næraa, thomas f. kokfelt, john c. schumacher and anders scherstén majaqqap qaavamajaqqap qaavamajaqqap qaava 63°15´n63°15´n63°15´n 10 km 63°15´n63°15´n63°15´n 50°30´w50°30´w50°30´w fiskenæssetfiskenæssetfiskenæsset 50°30´w50°30´w50°30´w undifferentiated ttg gneiss ilivertalik granite/charnockite fiskenæsset complex undifferentiated amphibolite/ dolerite dyke glacier quaternary grædefjo rd fi sk ef jo rd en mafic granulite greenlandfig. 1. simplified geological map of the central part of the north atlantic craton in southern west greenland showing majaqqap qaava in the fiskenæsset complex. based on maps published by the geological survey of denmark and greenland. © geus, 2010. geological survey of denmark and greenland bulletin 20, 67–70. open access: www.geus.dk/publications/bull 6868 the fiskenæsset complex (a. polat, personal communication 2010). among the dated, hand-picked grains there appear to be two populations, one at 2.92 ga and another at 2.87 ga (fig. 2a). a third possible component at c. 2.70 ga might represent a minor population of metamorphic grains. a known thermal event at 2.80 ga involving granulite-facies metamorphism that affected the western part of the fiskenæsset complex and the intrusion of the ilivertalik granite (t. næraa & a. scherstén, unpublished data) is not recorded among the dated zircon grains from majaqqap qaava. in situ observations to better understand the three zircon-forming events (c. 2.92, c. 2.87 and c. 2.7 ga), we observed the zircon grains in situ in polished slabs of the anorthosite samples using the scanning electron microscope at geus and the electron microprobe at the university of copenhagen. zircon grains occur in four different textural settings: (1) associated with ilmenite within the hornblenditic dyke (fig. 3a), (2) within the hornblenditic dyke, (3) in melt pockets associated with the hornblenditic dyke (fig. 3b) and (4) in cracks associated with chlorite (fig. 3c). a further feature in the hornblenditic dyke in the anorthosite is the break-down reaction of the ilmenite in the hornblende to form rutile, titanite and chlorite (fig. 3d). we dated zircon grains from these four different settings in situ using the same icp-ms instrument as described above. interpretation based on the in situ observations and measurements, our current understanding of the igneous and metamorphic history of the anorthosite at majaqqap qaava is as follows: after intrusion of the anorthosite at c. 2.97–2.95 ga, zircon formed at high-temperature conditions, e.g. from a reaction between baddeleyite and ilmenite. some of these zircon grains can be observed next to ilmenite grains (fig. 3a); however, no concordant ages were obtained from the in situ measurements. a later thermal event occurred at c. 2.92 ga, which forms the major age component in the zircon population extracted from the crushed sample (fig. 2a). since only one in situ zircon measurement yields 2.92 ga, the true nature of this event is unclear. this age might be related to an igneous event that formed the precursors to the amphibolite units in the area (see e.g. nutman et al. 2004), or the earliest onset of tonalitic gneiss formation in the region (næraa & scherstén, unpublished data). zircon grains from the contact region between the hornblenditic dyke and the anorthosite were observed both in relation to melt pockets (fig. 3b) and as occurring in the amphibole-anorthite assemblages. these different textural settings yield indistinguishable zircon 207pb/206pb ages of 2.878 ± 0.011 ga (mean square weighted deviation = 0.33) and 2.856 ± 0.016 ga (mean square weighted deviation = 1.3), respectively. there is, however, a large range in ages between the individual analyses. we suggest that this wide age range originates from inheritance or metamorphic overprinting, but the mean age reflects the crystallisation or resetting related to the intrusion of the hornblenditic dykes. the melt pockets in the hornblenditic dykes are likely to represent their final solidification. if correct, and if this age is representative, it implies that the dykes represent a late magmatic event, much later than previously assumed. this interpretation is at odds with field observations, which are best exb age ±2sd fraction ±2sd 2.720 28 0.04 0.06 2.872 5 0.56 0.22 2.919 7 0.40 -- relative misfit = 0.679 2950 2850 0.17 0.18 0.19 0.20 0.21 0.22 1.3 1.5 1.7 1.9 2.1 n u m b er o f an al ys es a 0 2 4 6 8 10 12 14 3.05 2.75 2.65 207 pb / 206 pb age (ga) 207 pb / 206 pb age (ga) 2.6 2.7 2.8 2.9 3.0 3.12.6 2.7 2.8 2.9 3.0 3.1 3.2 age ±2sd n mswd 2.697 65 3 5.9 2.863 7 30 3.0 0.17 0.18 0.19 0.20 0.21 0.22 1.5 1.7 1.9 2.1 238u/206pb 2.75 2.65 2.95 238u 206 pb 206 pb pb 207 fig. 2. zircon u/pb – pb/pb age distributions from hornblenditic dyke material and anorthosite from sample ggu 508216. a: zircon grains handpicked from crushed material with 90–110% concordant grains. unmixing of all grains following ludwig (2003). b: in situ dating of zircon grains with 90–110% concordant grains. mswd: mean square weighted deviation. n: number of analyses. the green lines show ages discussed in the text. 69 plained by their intrusion into a partly solidified anorthosite crystal mush rather than a brittle solid. alternatively, part of the hornblenditic dykes could have remelted during the intrusion of tonalitic gneisses in the area at this time (nutman et al. 2004; næraa & scherstén 2008). after further cooling hematite lamellae exsolved in the ilmenite, and these lamellae are seen as thin needles in the ilmenite grains (fig. 3d). the observed reaction microstructures (fig. 3d) suggest partial hydration of the assemblage hornblende + ilmenite, which yields the reaction products chlorite + rutile + titanite. this assemblage does not contain quartz. the reaction products are concentrated at the grain boundaries between hornblende and ilmenite, which suggests that the reactions are driven by small amounts of fluid present at the grain boundaries. since the anorthosites are dry, the extent to which retrograde metamorphic changes can be recorded is a function of the amount of water brought into the system by the hornblenditic dykes. the reactions are likely to have ceased after all the local fluid was consumed. assuming the reactions took place in an essentially closed system, as waterrich chlorite grew, the composition of the fluid could show considerable variation. within the chlorite-filled cracks in the hornblende-rich parts of the sample, newly grown zircon grains up to 100 μm in length are found (fig. 3c). these zircon grains appear to fill the interstitial space between the chlorite-rimmed hornblende grains. the source of zirconium to form these zircons is most probably the ilmenite grains that broke down in the reaction discussed above. the age of the zircon grains in these chlorite-filled cracks is poorly constrained at 2.70 ga (fig. 2b), but this age is in good accordance with the interpretation that the ilmenite break-down reaction occurred a c d bzircon zirconzircon zirconzircon pyritepyrite zircon hornblende ilm en ite quartz hornblende chlorite pyrite anorthite anorthite albite zircon pyrite chlorite ru til e hornblende ilmenite he m at ite titanite 200 μm 100 μm 250 μm 200 μm zircon hornblende chlorite ilmenite + titanite + rutile chlorite chlorite titanite rutile hornblende hornblende anorthite chloritechlorite fig. 3. backscatter electron contrast mode scanning electron microscope images showing the textural association of zircon grains in the sample ggu 508216 and the observed break down reaction of ilmenite. a: zircon associated with ilmenite. b: zircon grains in melt pockets associated with the hornblenditic dyke that intruded into the anorthosite. c: zircon in chlorite-filled cracks. d: ilmenite in hornblende has reacted with water to form chlorite, titanite and rutile. detail of the central grain shown in a. 7070 shortly after peak metamorphism (see below). regional metamorphism was previously dated at 2.72 ga, based on material from the nuuk region (e.g. friend et al. 1996) and the same age was reported north of ilivertalik by næraa & scherstén (2008). modelling of the metamorphic reaction modelling of reactions to determine the approximate conditions of formation is complicated by extensive compositional variation of the amphibole and by potential variation in the fluid composition. nevertheless, when using perplex (connolly 2005) it is possible to locate mineral composition isopleths that approximate microprobe data for the amphibole and chlorite. modelled compositions are: chlorite: xmg = 0.81–0.84; hornblende: xmg = 0.70–79; al per formula unit = 1.490–1.626, and measured compositions are: chlorite: xmg = 0.65–0.75; hornblende: xmg = 0.76–0.80; al per formula unit = 1.50–1.75. the perplex modelling results are shown in fig. 4. these results fit well with peak metamorphic conditions suggested by keulen et al. (2009). figure 4 shows the pressure–temperature region relevant to the mineralogy of the studied sample. the reaction seems to have occurred just after peak metamorphic conditions at about 600°c and after peak metamorphism in the area. as a result of this pilot study on zircon grains and their surrounding minerals in samples from majaqqap qaava within the fiskenæsset complex, southern west greenland, we are able to show that the anorthosite records a metamorphic history that is more complex than previously recognised. careful in situ observations prove helpful in unravelling the history of these rocks. acknowledgements alfons berger is thanked for help at the microprobe, and fiorella fabra aguilera and mojagan alaei are thanked for help with sample preparation. references connolly, j.a.d. 2005: computation of phase equilibria by linear programming: a tool for geodynamic modeling and its application to subduction zone decarbonation. earth and planetary science letters 236, 524–541. frei, d. & gerdes, a. 2009: precise and accurate in situ u–pb dating of zircon with high sample throughput by automated la-sf-icp-ms. chemical geology 261, 261–270. friend,  c.r.l., nutman,  a.p., baadsgaard,  h., kinny,  p.d. & mcgregor,  v.r. 1996: timing of late archaean terrane assembly, crustal thickening and granite emplacement in the nuuk region, southern west greenland. earth and planetary science letters 142, 353–365. keulen, n., scherstén, a., schumacher, j.c., næraa, t. & windley, b.f. 2009: geological observations in the southern west greenland basement from ameralik to frederikshåb isblink in 2008. geological survey of denmark and greenland bulletin 17, 49–52. ludwig, k.r. 2003: mathematical-statistical treatment of data and errors for 230 th/u geochronology. uranium-series geochemistry, reviews in mineralogy and geochemistry 52, 631–656. myers, j.s. 1985: stratigraphy and structure of the fiskenæsset complex, southern west greenland. bulletin grønlands geologiske undersøgelse 150, 72 pp. næraa, t. & scherstén, a. 2008: new zircon ages from the tasiusarsuaq terrane, southern west greenland. geological survey of denmark and greenland bulletin 15, 73–76. nutman, a.p., friend, c.r.l., barker, s.l.l. & mcgregor, v.r. 2004: inventory and assessment of palaeoarchaean gneiss terrains and detrital zircons in southern west greenland. precambrian research 135, 281–314. p u m p el ly it eac ti n o lit e fa ci es e p id o te am p h ib o lit e fa ci es greenschist facies sil ky and c h l r u t it h b l ilm calculated geotherms facies boundaries univariant reaction boundaries isopleths (constant composition) ss ss granulite facies ~2.70 ga ~3.20 ga ~2.56 ga ~3.00 ga hblss ilm h2o => ru tit chlss amphibolite facies 7 1 temperature (°c) 300 400 500 600 700 800 35 25 15 5 d ep th ( km ) 3 5 9 11 p re ss u re ( kb ar ) fig. 4. pressure–temperature (p–t) diagram that shows ranges of p–t estimates for modelled compositions of chlorite. the geotherms are based on measured and estimated content as a function of time of radiogenic elements in basaltic and felsic greenland rocks (see keulen et al. 2009 for further explanation). the pale-green area is the peak p–t conditions for regional metamorphism as suggested by keulen et al. (2009). the blue arrow is part of a possible cooling path. ru: rutile. tit: titanite. chlss: chlorite solid solution. hblss: hornblende solid solution. ilm: ilmenite. ky: kyanite. sil: sillimanite. and: andalusite. authors’ addresses n.k., t.n. & t.f.k., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ntk@geus.dk j.c.s., department of earth sciences, university of bristol, bristol bs8 1rj, uk. a.s., department of earth & ecosystem sciences, lund university, sölvegatan 12, s-223 62 lund, sweden. geological survey of denmark and greenland bulletin 17, 2009,61-64 the skaergaard intrusion (fig. 1) is probably the most studied layered gabbro intrusion in the world (wager & deer 1939; wager & brown 1968; mcbirney 1996; nielsen 2004). the intrusion is c. 54.5 ma old and was formed during the palaeogene opening of the north atlantic ocean, intruding into the base of the east greenland flood basalts. the intrusion is relatively small with a volume of c. 300 km3 (nielsen 2004). spectacular magmatic layering and systematic evolution in the compositions of liquidus phases and estimated melt compositions (e.g. wager & brown 1968) have made the intrusion the most studied example of the development of the ‘fenner trend’ of iron enrichment in basaltic liquids (e.g. thy et al. in press; veksler in press). the identification in the late 1980s of significant platinum-group elements (pge) and gold (au) occurrences in the intrusion (e.g. bird et al. 1991; nielsen et al. 2005) has led to continued investigation and exploration drilling. the skaergaard intrusion is suggested to hold c. 33 million ounces (1000 tonnes) of pge and c. 13 million ounces (400 tonnes) of au (nielsen et al. 2005). the mineralised zone is located in a c. 100 m thick zone of anomalous pge and au enrichment in the upper part of the middle zone (bird et al. 1991; nielsen et al. 2005) of the layered series. the mineralised zone consists of a succession of bowl-shaped, stratiform and very tightly controlled levels of palladium (pd) enrichment referred to as pd1 to pd5 (fig. 2; nielsen et al. 2005). the bottom level, pd5, is developed from margin to margin of the intrusion, whereas the overlying levels pd4 to pd1 are increasingly restricted in width, and the entire succession of pd levels is only developed in the central part of the intrusion. the structure of the mineralised zone can be compared to a set of bowls with upward-decreasing diameters. gold is always concentrated in the uppermost palladium levels or in a level above the top palladium level, irrespective of the number of developed pd levels. more detailed descriptions are provided by nielsen et al. (2005). the exploration drill cores provide material and structural information from previously inaccessible parts of the intrusion (nielsen et al. 2005). the 3-d image presented in fig. 4 is based on drill-core information (petrographical, petrophysical, geochemical etc.) and surface information. it allows an unprecedented insight into the internal structure of the upper part of the intrusion and offers a possibility of refinement of volume estimates and quantitative modelling of the developing a 3-d model for the skaergaard intrusion in east greenland: constraints on structure, mineralisation and petrogenetic models troels f.d. nielsen, símun d. olsen and bo m. stensgaard © geus, 2009. geological survey of denmark and greenland bulletin 17, 61–64. available at: www.geus.dk/publications/bull 61 fig. 2. characteristic variation in whole-rock pd concentration in the central parts of the skaergaard mineralised zone (core ddh 90-22, from bernstein & nielsen 2004). fig. 1. the skaergaard intrusion is located in the kangerlussuaq region in the palaeogene magmatic province in east greenland. kangerlussuaq iceland inland ice 100 km skaergaard intrusion blosse vil le k ys t tasiilaq illoqqortoormiut /scoresbysund palaeogene intrusive centre pre-cretaceous sedimentary rocks cretaceous–palaeogene sed. rocks palaeogene basalt greenland 69°n 66° 32°w 24° d ep th in c or e d d h 9 022 ( m ) 0 1000 2000 3000 concentration (ppb) gold palladium pd5 pd4 pd3 pd2 pd1 pd1/au au+1980 990 1000 1010 1020 1030 1040 rosa_2008:rosa-2008 01/07/09 15:48 side 61 zones and subzones of the intrusion. a constrained structural model will in turn allow evaluation and revision of crystallisation models for the basaltic liquid in the magma chamber. 3-d modelling of the intrusion and the mineralised zone the initial aim of the 3-d modelling was a visualisation of the intrusion and the associated pge and au mineral occurrences. geographical information system software was used for the compilation of the surface data used for the model. these data, together with subsurface data, were subsequently imported into modern 3-d mining and resource software (gemcom gems®), which was used for the construction of the 3-d model of the intrusion and its mineralised zone. the detailed topographical model needed for the modelling (fig. 3) was constructed from satellite aster data (resolution 30 × 30 m). aster scenes and the 1:20 000 scale geol ogical map of the intrusion and adjacent area were draped on the terrain model. forty-one cores with a total length of 23 425 m have been drilled since 1989. the deepest holes reached levels of c. 1200 m below the collars of the drill holes. the petrographic variation in all these cores is described in drill-hole logs in company reports in the archives of the geological survey of denmark and greenland. these logs were digitalised and compiled. the courses of the drill holes (taking azimuth and dip into account) were visualised in 3-d, and assays for pge and au displayed together with the petrographic information. all the information was subsequently assessed for each drill hole, and the delineation of specific lithologies and mineralised sections was interpolated manually by the geologist software operator from one drill hole to another and from drill holes to surface exposures of the mineralised zone. triangulation surfaces were constructed mathematically by the gems software from the delineations and united into wire-frames that represent 3-d solids (geological bodies). the delineation and resulting solids were validated by the software. dykes in the intrusion were also modelled as solids. mapped-out fault planes were visualised as 3-d surfaces. in intrusion-wide images the mineralised zone is a very narrow structure. the 3-d model is best seen ‘live’, and we have chosen, as examples, to show the initial results of the imaging of the mineralised zone in two vertical 2-d panels through the intrusion (figs 4, 5). in fig. 5 the mineralised zone is shown as the zone between the lower boundary of the lowermost pd-levels (pd5, cut-off at c. 1 gram per tonne pd) and the top of the au-rich part of the mineralised zone (pd1/au or au + 1 levels, cut-off at c. 0.8 gram per tonne au). 62 fig. 3. aster satellite image with topography (1.5 × vertical exaggeration; see text for explanation). the green line shows the outer boundary of the skaergaard intrusion, and the blue dots show the locations of the collars of drill holes. 2 km mbs uz southsouth easteast northnorth westwest south ubs east mbs uznorth mzlz mbs west fig. 4. satellite image showing the area of the mineralised zone so far covered by the 3-d model (purple area) with the geological boundaries transferred from the geological map of the skaergaard intrusion (mcbirney 1989). mbs, marginal border series (wall rocks); ubs, upper border series (roof rocks); layered series (ls; floor rocks) consisting of lz, lower zone; mz, middle zone; uz, upper zone. islands in ubs: the younger basistoppen sill. the two white lines show the location of the 2-d cross-sections (fig. 5). abbreviations also apply to table 1 and fig. 6. rosa_2008:rosa-2008 01/07/09 15:48 side 62 results the west –east section of fig. 5a shows the mineralised zone to be bowl-shaped with a central depression of c. 400 m. the magnitude of the depression is in broad agreement with the margin-to-margin depression of c. 700 m modelled by nielsen (2004) and nielsen et al. (2005). the difference in the magnitude of the depression reflects that the 3-d model does not reach all the way to the margins of the intrusion. as expected, the imaging also shows that the vertical distance between the lower and upper boundaries of the mineralised zone increases towards the centre of the intrusion, in agreement with the structure of the mineralised zone proposed by nielsen et al. (2005). the demonstrated bowl-shape of the mineralised zone, and thus the layered gabbros, corroborates the model suggesting concentric crystallisation of the gabbro on the floor, walls and below the roof of the intrusion (nielsen 2004). the north–south section (fig. 5b) shows the general 20° dip of the layered gabbros and the mineralised zone. application of 3-d modelling to the evolution of the skaergaard intrusion nielsen (2004) developed a structural model for the intrusion solely on the basis of field observations and analogies. compared to the classic and traditional accumulation models, the apparent concentric crystallisation in the 300 km3 magma chamber reduces the volumes of the most evolved zones and subzones in the intrusion and thus the proportions of the products of the crystallisation process. the model was used for a mass balance-based estimate of the bulk composition of the intrusion, which turned out to be a composition very similar to that of the contemporaneous flood basalts. the data in nielsen (2004) can also be used for the calculation of the line of liquid descent (lld) of the bulk liquid (fig. 6; table 1). toplis & carroll (1995) modelled the lld for the skaergaard intrusion on the basis of experimental investigations. as shown in fig. 6 the trend of the supposed skaergaard liquid of toplis & carroll (1995) has the same shape as the one calculated using mass balance (table 1). the data in fig. 6 are projected from the sio2 corner above the plane in the figure, and the difference between the toplis & carroll (1995) and the lld suggested here is a reflection of differences in the starting compositions. 63 west eastsea level terrain surface mineralised zone north southsea level mineralised zone terrain surface 1 km 1 km b a fig. 5. west–east (a) and north–south (b) cross-sections through the skaer gaard intrusion. the blue and green lines show the lower and upper boundaries of the skaergaard mineralised zone (see text for definition). the bowl-shape and the increasing stratigraphic width of the mineralised zone towards the centre of the intrusion are seen (see also nielsen et al. 2005). table 1. compositions of liquids during the fractionation of the skaergaard magma bulk l1 l2 l3 l4 l5 l6 l7 sk-tfdn lzb lzc mz uza uzb uzc mg % solidified 0.00 32.74 55.26 62.56 75.81 85.57 93.43 95.00 sio 2 47.88 47.03 45.92 46.52 48.19 50.73 58.45 62.08 tio 2 3.03 3.74 4.86 4.63 3.74 2.65 1.40 1.07 al 2 o 3 13.87 11.93 11.24 11.15 10.75 10.42 11.35 12.39 fe 2 o 3 2.00 2.34 2.70 2.72 2.80 2.78 2.12 1.63 feo 13.32 15.63 17.97 18.12 18.64 18.53 14.13 10.88 mno 0.22 0.26 0.29 0.30 0.32 0.34 0.29 0.20 mgo 6.29 6.18 4.62 4.15 3.16 1.83 0.56 0.63 cao 10.16 9.62 8.78 8.55 7.89 7.37 5.89 4.64 na 2 o 2.56 2.48 2.61 2.71 2.94 3.18 3.78 4.18 k 2 o 0.40 0.44 0.55 0.61 0.79 1.07 1.63 1.99 p 2 o 5 0.27 0.35 0.47 0.54 0.78 1.10 0.42 0.31 sum 100.00 100.00 100.01 100.00 100.00 100.00 100.02 100.00 mg no. 0.457 0.413 0.314 0.290 0.232 0.150 0.066 0.094 based on bulk liquid sk-tfdn in nielsen (2004). the composition of the liquid as it evolves is calculated by subtraction of average compositions of correlated ls, mbs and ubs subzones (mcbirney 1989) in the mass proportions in nielsen (2004). the spread sheet with the calculation is available on request. the composition of the liquid of a specific zone refers to the composition of the liquid at the base of the indicated zone. bulk composition is corrected so that the end-result matches the composition of average melanogranophyre (mg). fe 2 o 3 /feo has been set at 0.15. the abbreviations are explained in fig. 4. rosa_2008:rosa-2008 01/07/09 15:48 side 63 64 the lld of the skaergaard intrusion is of utmost scientific interest. well-constrained deviations from the expected can be reflections of processes that have not been taken into account in the modelling of the fractionation process. the lack of balance in the sio2 distribution, as reflected in the common quartz-normative compositions of the upper bor der series of the intrusion (naslund 1984), was suggested to reflect chemical stratification of the cooling magma (hoover 1989). but what process would have been responsible for the chemical stratification? sio2-enrichment in the upper part of the magma chamber could be due to liquid immiscibility (e.g. jakobsen et al. 2005) or dynamic conditions. only with well-constrained mass balances and geophysical models for the shape of the magma chamber can numeric models for the evolution of the skaergaard intrusions be developed and the relative importance of all the suggested processes in the evolution of the melt evaluated. all well-constrained internal boundaries and the details of the mineralised zone (bulk chemistry, lithologies and mine ralogy) in the skaergaard intrusion will be included in the 3-d model in the coming years. this will allow refinement of the 3-d distributions and volumes of different lithologies, including the mineralised zone, lead to more advanced massbalance models for the skaergaard intrusion, and provide more general constraints for modelling of the crystallisation and fractionation processes in basaltic magma chambers. references bernstein, s. & nielsen, t.f.d. 2004: chemical stratigraphy in the skaer gaard intrusion. danmarks og grønlands geologiske under søgelse rapport 2004/123, 31 pp. + 1 cd. bird, d.k., brooks, c.k., gannicott, r.a. & turner, p.a. 1991: a goldbearing horizon in the skaergaard intrusion, east greenland. econo mic geology 86, 1083–1092. hoover, j.d. 1989: petrology of the marginal border series of the skaer gaard intrusion. journal of petrology 30, 399–439. jakobsen, j.k., veksler, i.v., tegner, c. & brooks, c.k. 2005: immiscible ironand silica-rich melts in basalt petrogenesis documented in the skaer gaard intrusion. geology 33, 885–888. mcbirney, a.r. 1989: geological map of the skaergaard intrusion, east greenland. eugene, usa: university of oregon. mcbirney, a.r. 1996: the skaergaard intrusion. in: cawthorn, r.g. (ed.): layered intrusions, 147–180. amsterdam: elsevier. mcbirney, a.r. & naslund, h.r. 1990: the differentiation of the skaer gaard intrusion. a discussion of hunter and sparks (contributions to mineralogy and petrology 95, 451–461). contributions to mineralogy and petrology 104, 235–240. naslund, h.r. 1984: petrology of the upper border series of the skaer gaard intrusion, east greenland. journal of petrology 25, 185–212. nielsen, t.f.d. 2004: the shape and volume of the skaergaard intrusion, greenland: implications for mass balance and bulk composition. journal of petrology 45, 507–530. nielsen, t.f.d., andersen, j.c.ø & brooks, c.k. 2005: the platinova reef of the skaergaard intrusion. in: mungal, j.e. (ed.): exploration for platinum group element deposits. mineralogical association of canada short course series 35, 431–455. thy, p., lesher, c.e. & tegner, c. in press: the skaergaard liquid line of descent revisited. contributions to mineralogy and petrology. toplis, m.j. & carroll, m.r. 1995: an experimental study of the influence of oxygen fugacity on fe-ti oxide stability, phase relations, and mineral–melt equilibria in ferro-basaltic systems. journal of petrology 36, 1137–1170. veksler, i.v. in press: extreme iron enrichment and liquid immiscibility in mafic intrusions: experimental evidence revisited. lithos. wager, l.r. & brown, g.m. 1968: layered igneous rocks, 588 pp. edin burgh: oliver & boyd. wager, l.r. & deer, w.a. 1939: geological investigations in east green land, part iii. the petrology of the skaergaard intrusion, kanger dlugssuaq, east greenland. meddelelser om grønland 105(4), 352 pp. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: tfn@geus.dk wo + en + fs uzpy ro xe ne (s) plagioclase ab + or an oxygen units [ol, qz] mg# = 0. 05 lz mz mg# = 0. 5mg# = 0. 25 mg# = 1 fig. 6. skaergaard liquid lines of descent and immiscible liquids in the (ab + or)–an–(wo + en + fs) projection from veksler (in press), who used the lld compositions shown in table 1. black dots: model liquids in the same sub-zones of the layered series calculated by mass balance (table 1). circles: trapped skaergaard liquids (mcbirney & naslund 1990). thick dashed curve: experimental line of liquid descent (toplis & caroll 1995). the dashed lines show the boundary between plagioclase and pyroxene crystallisation fields at different mg numbers (mg/mg + fe; see veksler in press for details). rosa_2008:rosa-2008 01/07/09 15:48 side 64 geological survey of denmark and greenland bulletin 20, 2010, 51–54 51 the kingdom of denmark (denmark, faroe islands and greenland) ratified the united nations convention on the law of the sea (unclos) on 16 november, 2004 and this allows for a period of ten years to submit extended continental shelf claims beyond 200 nautical miles (nm) to the commission on the limits of the continental shelf. to acquire the necessary data for delineating the extended continental shelf, the continental shelf project of the kingdom of denmark was launched by the ministry of science, technology and innovation in cooperation with the faroese and greenland governments. several institutions participate in the project. the technical work for the greenland part is coordinated by the geological survey of denmark and greenland (geus) and the coordination of the faroese part is shared between the faroese earth and energy directorate (jarðfeingi) and geus. further information can be found on the project website www.a76.dk. background article 76 of unclos is the key to future jurisdiction over resources on and below the seabed beyond the 200 nm limit. the right to explore and exploit these resources, which include both non-living resources (hydrocarbons and minerals) and bottom-dwelling living resources, may have significant economic implications. according to article 76, a variety of scientific and technical data are required to be submitted to the commission on the limits of the continental shelf. these include geodetic, bathymetric, geophysical and geological data regarding e.g. the 200 nm and 350 nm limits from the territorial baselines, the location of the foot of the continental slope, the 2500 m isobath and the sediment thickness beyond the foot of the slope. the latter is defined as the point of maximum change in the gradient at the base of the continental slope. areas of interest – the faroe islands the continental shelf of the faroe islands extends beyond 200 nm in two geographical regions, here referred to as the northern and the southern area (figs 1, 2). the geomorphological and geological settings are fundamentally different in the two areas. already in the early phase of the project, informal arrangements were made with the neighbouring states for exchange of data, and these data form a significant part of the overall data base in both areas. field work has been carried out in 2003, 2004, 2005, 2007 and 2008. a wide range of data and samples has been collected including seismic reflection and refraction data, singleand multibeam echo sounder bathymetric data, sediment cores, marine gravity data and airborne magnetic data. northern area. north of the faroe islands (fig. 1), the continental margin is characterised by a number of ridges and elevated sea-floor highs that extend from the shelf and slope region into the northern deep. the most pronounced of these sea-floor highs is the ægir ridge, which is an extinct part of the spreading system that created the oceanic sea floor of the northern deep. the northern deep is a large sedimentary the continental shelf project of the kingdom of denmark – status at the beginning of 2010 christian marcussen and martin v. heinesen æ g i r r i d g e møre margin 150 km jan m ayen fracture z o ne faroe islands northern deep vøring plateau n orth sea fan fa ro esh et la nd c ha nn el faroe-iceland ridge 62°n 66°n 12°w 4°w 4°e fig. 1. the northern continental shelf of the faroe islands. the continental shelf beyond 200 nm in the area north of the faroe islands, as delineated in the submission, amounts to 87 792 km2 and is highlighted. green: agreed borders within 200 nm. red: faroese 200 nm limit. yellow: iceland’s and norway’s eez (exclusive economic zone; 200 nm). orange: outer limit of the continental shelf. © geus, 2010. geological survey of denmark and greenland bulletin 20, 51–54. open access: www.geus.dk/publications/bull 5252 basin with sediment thicknesses locally reaching 3 km. the continental margin is strongly affected by slope-related processes (mass wasting, turbidity currents, etc.) and complex erosional and depositional actions of strong ocean currents. after a period of processing, interpreting and analysing the data in accordance with article 76 of unclos and the technical and scientific guidelines of the commission on the limits of the continental shelf (1999), a special task force was established with the responsibility of producing the formal documentation for the northern area to be submitted to the commission. nine task force members were selected for their expertise in a wide range of geoscientific and legal disciplines and they represent a number of governmental institutions and ministries in denmark and the faroe islands. the task force initiated its preparation of the submission in april 2008, and by april 2009 the documentation was delivered to the danish ministry of foreign affairs who formally submitted it to the commission on the limits of the continental shelf on 29 april, 2009 on behalf of the kingdom of denmark. on 27 august, 2009 the submission was presented to the commission by a danish–faroese delegation during its 24th formal session. the continental shelf beyond 200 nm in the area north of the faroe islands, as delineated in the submission, amounts to 87 792 km2 (fig. 1). the area overlaps with regions that are also covered by submissions of the two neighbouring states, norway and iceland. the three nations have made a mutual agreement on how to divide the continental shelf area between themselves that is without prejudice to the work of the commission on the limits of the continental shelf. southern area. the southern area (fig. 2) is dominated by the continental faroe-rockall plateau that extends several hundred nautical miles toward the south-west. three additional states, the united kingdom, ireland and iceland, include the area as part of their outer continental shelf. geologically and geomorphologically the plateau is characterised by a number of banks and basins where continental crust and older sediments are covered by palaeogene synand post-breakup volcanic strata that are similar to those found on the faroe islands. the marginal parts of the plateau are characterised by extensive sediment accumulations, mainly contourites that were deposited by bottom currents along the lower part of the continental slope. the task force is presently preparing the submission documents for the southern area and expects to deliver the final documentation to the ministry of foreign affairs by the end of 2010. areas of interest – greenland three areas around greenland have been defined as areas of interest regarding an extension of the continental shelf beyond 200 nm: an area south of greenland (including the eirik ridge), an area off north-east greenland near the east greenland ridge and an area north of greenland along the lomonosov ridge and in the amundsen basin (figs 3–5). the three areas are briefly described in marcussen et al. (2004). as with the faroe islands work commercial contractors were used to acquire seismic reflection and bathymetric data. however, because of the sea-ice conditions off northeast and north greenland data from these regions have been acquired in cooperation with canada and sweden. area south of greenland. the region south of greenland (fig. 3) is characterised by a fairly narrow shelf and by thick sedimentary successions between greenland and canada, especially within the extinct spreading zone. the main tasks 16°w 250 km 8°w 0°w 60°n 56°n 64 °n feni d rift faroe islands faro e-sh etlan d c h an n el cha rlie g ibb s fr ac tur e z one e d o ras b an k h atto n b an k faroe island ridge icelan d b asin h atto n r o ckall b asin r o ckall b an k r o ckall t h ro u gh eriador seamount porcupine b ank f a r o e r o c k a l l p l a t e a u fig. 2. the southern continental shelf of the faroe islands. green: agreed borders within 200 nm. red: faroese 200 nm limit. yellow: iceland’s, ireland’s and uk’s eez (200 nm). the outer continental shelf in this area amounts to c. 500 000 km2. 53 in this region are threefold: to map the bathymetry of the continental slope especially along the eirik ridge in order to outline the foot of the continental slope, to map the sediment thickness and to investigate the nature of the eirik ridge. two seismic surveys were carried out in 2003 and 2006 to map the sediment thickness. a total of c. 4000 km of data were acquired. other already existing commercial and scientific data supplement these acquired data. in 2008 the bathymetry of the continental slope and parts of eirik ridge was mapped by multibeam echo sounding. this data set was supplemented in 2009 with data acquired during a german research cruise (uenzelmann-neben 2009). also in 2009, a joint canadian-danish cruise mapped the deeper structure of the eirik ridge using refraction seismic profiling (funck et al. 2009). data acquisition in the area south of greenland within the continental shelf project is now completed and a task force will prepare the submission documents for this area in 2011. area north-east of greenland. this region (fig. 4) is characterised by a wide continental shelf. the east greenland ridge is assumed to be a natural prolongation of north-eastern greenland. in the summer of 2002, geus and the university of bergen carried out a joint seismic refraction and reflection survey of the east greenland ridge. interpretations of the refraction data show that the velocity structure of the east greenland ridge is consistent with continental crust (døssing et al. 2008). during the lomrog i cruise with the swedish icebreaker oden bathymetric mapping of the east greenland ridge and the continental slope of the east greenland shelf south-west of the ridge was initiated. the plan is to complete this mapping over the next few years and to supplement it with seismic data and geological sampling. area north of greenland. the lomonosov ridge and the morris jesup rise are assumed to be natural prolongations of northern greenland (fig. 5). older seismic data from the amundsen basin have shown sediments with sufficient thickness to be used in extending the continental shelf (weigelt & jokat 2001). due to the difficult year-round sea-ice conditions in this region and the lack of appropriate danish logistical possibilities and research platforms, data collection in this area requires cooperation with other countries (macdougall et al. 2008). therefore, in 2005 a memorandum of understanding was signed with canada that forms the basis for six major data acquisition programmes mainly in the area north of greenland. in the spring of 2006, seismic refraction data were collected from the sea ice during the lorita expedition (jackson & dahl-jensen 2007). field work in the spring of 2009 concentrated on bathymetric and gravimetric data acquisition supplemented by aerogeophysical measurements. in order to acquire seismic data under difficult ice conditions, a seismic reflection acquisition system has been developed in cooperation with the department of earth sciences, aarhus university, based on experience gained from other surveys in the arctic by canadian, german, norwegian and us institutions. by cooperating with the swedish polar research secretariat it was possible to use oden during two cruises (jakobsson et greenland eirik ridge 250 km canada 60ºw 51ºw 42ºw 54ºn 60ºn fig. 3. the potential claim area off southern greenland (grey tone) has a size of approximately 45 000 km2. green: agreed border within 200 nm. red: greenland’s eez. yellow: canada’s eez. blue: 350 nm limit. dashed line: unofficial median line between greenland and canada. 80ºn 72ºn 20ºw 0º 20ºe 250 km svalbard east greenland ridge jan mayen g re en la nd fig. 4. the potential claim area off north-east greenland (grey tone) with an approximate size of 63 000 km2. green: agreed boundaries within the 200 nm limit between greenland–svalbard and greenland–jan mayen. red: greenland’s eez. yellow: norway’s eez from svalbard and jan mayen. blue: 350 nm limit. 5454 al. 2008, marcussen et al. 2010). further work is planned for either 2011 or 2012. since 2007 the canadian, russian and danish continental shelf projects have discussed scientific issues regarding the lomonosov ridge on a regular basis. conclusions since the start of the project the work has progressed according to plan and submissions for the two faroese areas will be completed in 2010. data acquisition is complete for the area south of greenland. due to difficult ice conditions, work in the areas north-east and north of greenland will not be completed until the end of 2014. due to the high number of submissions received by the commission on the limits of the continental shelf, final processing of the two partial submissions for the faroe islands may not start until several years after the submissions. the commission on the limits of the continental shelf could spend 1–2 years considering the documents before reaching a conclusion and submitting a recommendation for the final outer limits of the continental shelf. references commission on the limits of the continental shelf 1999: scientific and technical guidelines of the commission, clcs/11 (http:// www.un.org/depts/los/clcs_new/commission_documents.htm# guidelines). døssing, a., dahl-jensen, t., thybo, h., mjelde, r. & nishimura, y. 2008: east greenland ridge in the north atlantic ocean: an integrated geophysical study of a continental sliver in a boundary transform fault setting. journal of geophysical research 113, b10107, 33 pp. funck, t., dehler, s.a., chapman, c.b., delescluse, m., iuliucci, j., iuliucci, r., judge, w., meslin, p. & ruhnau, m. 2009: cruise report of the signal 2009 refraction seismic cruise (hudson 2009-019). danmarks og grønlands geologiske undersøgelse rapport 2009/74, 60 pp., 2 appendices + 1 cd. jackson, h.r. & dahl-jensen, t. (eds) 2007: field report for lorita 2006 – lomonosov ridge test of appurtenance. danmarks og grønlands geologiske undersøgelse rapport 2007/5, 184 pp. jakobsson, m., marcussen, c. & lomrog scientific party 2008: lomonosov ridge off greenland 2007 (lomrog) cruise report, 122 pp. copenhagen: geological survey of denmark and greenland. macdougall, j.r., verhoef, j., sanford, w. & marcussen, c. 2008: challenges of collecting data for article 76 in ice covered waters of the arctic. 5th ablos conference difficulties in implementing the provisions of unclos 15–17 october 2008, 21 pp. (http://www.gmat.unsw.edu. au/ablos/ablos08folder/session4-paper1-macdougall.pdf). marcussen, c., christiansen, f.g., dahl-jensen, t., heinesen, m., lomholt, s., møller, j.j. & sørensen, k. 2004: exploring for extended continental shelf claims off greenland and the faroe islands – geological perspectives. geological survey of denmark and greenland bulletin 4, 61–64. marcussen, c. & lomrog ii scientific party 2010: lomrog ii – continued data acquisition in the area north of greenland. 2009 yearbook of the swedish polar research secretariat, stockholm, 43–51. (http:// www.polar.se/organisation/pdf/yearbook2009.pdf). memorandum of understanding between the earth sciences sector of the department of natural resources of canada and the geological survey of denmark and greenland of the ministry of the environment of denmark concerning cooperation relating to the delineation of their continental shelves, signed 21 june 2005. the government of the kingdom of denmark together with the government of the faroes 2009: the continental shelf north of the faroe islands. partial submission. executive summary, 20 pp. (http://a76.dk/ xpdf/dnk2009executivesummary_s.pdf). uenzelmann-neben, g. (ed.) 2009: the expedition of the research vessel ‘maria s. merian’ to the labrador sea in 2009 (msm 12/2) reykjavik – reykjavik 17 june – 13 july 2009. berichte zur polarund meeresforschung 599, 91 pp. (http://hdl.handle.net/10013/epic.33539). weigelt, e. & jokat, w. 2001: peculiarities of roughness and thickness of oceanic crust in the eurasian basin, arctic ocean. geophysical journal international 145, 505–516. 0°80°n 85°n 30°w 30°e greenland svalbard canada lomonosov ridge morris jesup rise a m un ds en b as in 250 km fig. 5. the potential claim area off north greenland (grey tone) can have a size up to 150 000 km2. red: greenland’s eez. yellow: canada’s, norway’s and the russian federation’s eez. blue: the 350 nm limit. green: the agreed boundaries within 200 nm between greenland–svalbard and greenland–canada. stippled green line: greenland’s eez north of the agreed boundary with canada (1973, but within 200 nm). black stippled lines: unofficial median lines. authors’ addresses c.m., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: cma@geus.dk m.v.h., faroese earth and energy directorate (jarðfeingi), brekkutún 1, 188 hoyvík, faroe islands. geological survey of denmark and greenland bulletin 35, 2016, 59-62 59© 2016 geus. geological survey of denmark and greenland bulletin 35, 59–62. open access: www.geus.dk/publications/bull for more than 80 years the skaergaard intrusion, 68°n in southern east greenland, has been a foremost natural laboratory for the study of the crystallisation and fractionation of basaltic magma. this process has been of prime importance in the evolution of the earth and other stony planets. models that have been developed and refined during numerous studies of this particular intrusion have been part of the foundation for petrogenetic modelling for decades. in later years, vast amounts of new data have been added, due to systematic sampling in the field and from analysis of exploration drill cores. methods for the study on grain-size scale have advanced, and the quest for a wellsupported genetic model for the pge-au mineralisation of the intrusion has intensified. the new data and insight question the applicability of conventional petrogenetic modelling, and as a consequence, increasing importance is placed on in situ crystallisation and fractionation in mush zones at the roof, walls and floor of the intrusion. the skaergaard intrusion the skaergaard intrusion (wager & brown 1968) is a comparatively small but well-preserved and well-exposed layered gabbro intrusion (fig. 1a). it is 56 ma old (wotzlaw et al. 2012) and was emplaced during the opening of the north atlantic. it is 7 × 11 km in surface exposure, has a total structural height of c. 4 km, and, dependent on the chosen modelling paradigm, has a box-like (nielsen 2004) or ellipsoid shape (irvine et al. 1998, svennevig & guarnieri 2012) and a volume of c. 300 km3. the intrusion crystallised concentrically inward from the margins (fig. 1b) with the layered series (ls, lz, mz and uz) in the bowl-shaped floor, the marginal border series (mbs) on the walls, and upper border series (ubs) below the roof. the ubs and ls meet at the sandwich horizon (sh). all three series are subdivided on the basis of a parallel evolution in liquidus parageneses (salmonsen & tegner 2013 and references therein). new research initiatives petrogenetic modelling of the skaergaard gabbros and the evolution of the melt in the intrusion have traditionally rested on the textural interpretation of the gabbros as rocks composed of liquidus crystals, continued growth of these in situ fractionation and inward migration of the solidification front in the skaergaard intrusion, east greenland troels f.d. nielsen enewsw sh sh no vertical exaggeration 2 km mbsmbs sea level uzc uzb uza mz lzc lzb lza+hz ubsubs mbsmbs fault upper zone greenland 5 km miki fjord watkins fjord a b 31°25´w 68°15´n sediment other gabbros skaergaard intrusion flood basalts basement gneiss ubs uz mzmbs lz upper border series middle zonemarginal border series lower zonelayered series l ay er ed s er ie s ls fig. 1. a: geological setting of the skaergaard intrusion between basement gneisses, sediments, other gabbros and flood basalts. location in greenland in insert. b: reconstructed cross-section (nielsen et al. 2015). the upper and lower zones of the layered series are subdivided on the basis of liquidus paragenesis (see text). sea level and present topography are shown in black lines. 6060 crystals in equilibrium with the bulk liquid (adcumulus growth), and crystallisation of solids from trapped liquids. sorting of crystals on the magma chamber floor has been likened to processes established for clastic sediments, including stratification of crystal mushes in matrix-supported mass flows. despite challenges these models have remained robust, and most researchers are faithful to this classic cumulus paradigm and the modelling tools developed therefrom. research initiatives in the later decades of the 20th century (mcbirney 1996 and references therein) provided much new information and were accompanied by investigations facilitated by the exploration of pge-au mineralisation in the intrusion (e.g. bird et al. 1991). notable outcomes of this research include the development of double diffusive convection models (mcbirney & noyes 1979), evaluation of the petrogenetic importance of immiscibility between fe-rich and si-rich silicate melts (jakobsen et al. 2011 and references therein), as well as evidence for isotopic disequilibrium and tight age controls on the emplacement and solidification of the intrusion (see wotzlaw et al. 2012, and references therein). the studies of the skaergaard intrusion surged in 2000 with access to assay data and up to 1200 m long drill cores. the data in the public domain allowed calibration of structural models for the interior of the intrusion (nielsen 2004), erection of compaction models (tegner et al. 2009; mckenzie 2011), and studies of the mineralisation (nielsen et al. 2005, 2015; andersen 2006; rudashevsky et al. 2015 and references therein; holwell et al. 2015; keays & tegner 2015). petrographic studies focused, e.g. on compositional variations in plagioclase (namur et al. 2014) and on clinopyroxene-filled dihedral angles between plagioclase crystals (e.g. holness 2015 and references therein). changes in dihedral angles give indications for, e.g. the arrival of new phases on the liquidus of the silicate melts and changes in permeability, and thus for the controls on the mobility of elements of economic interest. c o re -m an tl e tr an si ti o n m an tl eri m t ra n si ti o n core mantle rim (b) 40 45 50 55 a n -c o n te n t (% ) 0 50 100 150 200 250 300 350 distance (μm) fig. 2. compositional variation in lzb plagioclase crystal with small core (liquidus) and broad mantle crystallised during in situ fractionation involving mush melt, and rim crystallised during buffered crystallisation (after namur et al. 2014). a plag plag s s s ol ol ol s s s s plag cpx cpx mt ilm ilm cpx feti b s fig. 3. a: transmitted light image (c. 13 × 9 mm) showing interconnected magnetite and ilmenite that crystallised from interstitial mush melt. b: electron microprobe backscatter image (see nielsen et al. 2015). symplectites (s) formed by reaction between plagioclase and reactive fe-rich mush melt that crystallised most of the magnetite and ilmenite in the view. cpx: clinopyroxene (with exsolutions in fig. 3b). feti: feti-oxides. ilm: ilmenite. mt: magnetite. plag: plagioclase. ol: olivine. s: symplectitic intergrowths. scale at base of the image. 61 the importance of in situ fractionation modelling based on the classic cumulate paradigm suggests that the proportion of trapped liquid decreased from 30– 50 per cent to only a few per cent during the solidification of the intrusion (tegner et al. 2009). this is, however, in conflict with reactions between liquidus minerals and ferich silicate melts (holness et al. 2011) and the occurrence of immiscible melt droplets throughout much of the floor cumulates (jakobsen et al. 2011 and references therein). they are supposed to result from extended in situ crystallisation and fractionation (langmuir 1989) in crystal mush, long residence time, and ineffective compaction. this is supported by the very common zonation in plagioclase (an example is shown in fig. 2, namur et al. 2014) and the distribution of magnetite crystallised from interstitial melt (fig. 3a; nielsen et al. 2015). toward a new solidification model a magma chamber will always be hot in the middle and crystallisation will always occur in the crystal mush between solidified gabbro and the remaining melt, unless the system is affected by vigorous convection. in the skaergaard intrusion this seems unlikely due to the concentric solidification (nielsen 2004). all gabbro samples have recorded a temperature interval on the line of liquid descent and all have witnessed the inward migration of the crystallisation front, fronts with new phases on the liquidus of the mush liquid, and the solidification front. the mushy layer is a sub-chamber of crystal mush migrating inwards, and the samples we collect reflect only processes within the mush itself and the bulk composition of the liquid that was processed in the mush (fig. 4). any sample of the gabbros is composed of minerals left behind by the inward-migrating mush layer. in broad terms, the composition of the floor gabbro is equal to bulk liquid minus what rose out of the floor, e.g., low density melt, and that of the roof gabbro is equal to what remained under the roof, e.g. low density minerals and melt (salmonsen & tegner 2013; nielsen et al. 2015). roof and floor series are complementary, and neither series represents the evolution of the bulk magma, but their weighted average does. the modelling of the evolution of layered intrusions is commonly only based on exposed floor cumulates, and the common neglect of complementary successions in the lost roof of the intrusions may therefore lead to erroneous petrogenetic conclusions. undoubtedly, future research in the skaergaard intrusion and its mineralisation will focus on very detailed petrography, mineralogy, in situ mineral chemistry and isotope geochemistry, and on unravelling of the complexities of the solidification processes. petrogenetic modelling on the basis of bulk rock chemistry without detailed petrographic information is prone to lead to significant oversimplification and unwarranted confirmation of the chosen models. references andersen, j.c.ø. 2006: postmagmatic sulphur loss in the skaergaard intrusion: implications for the formation of the platinova reef. lithos 92, 198–221. bird, d.k., brooks, c.k., gannicott, r.a. & turner, p.a. 1991: a goldbearing horizon in the skaergaard intrusion, east greenland. economic geology 86, 1083–1092. holness, m.b. 2015: plagioclase growth rates control three-grain junction geometry in dolerites and gabbros. journal of petrology 56(11), 2117–2144. holness, m.b., stripp, g., humphreys, m.c.s., veksler, i.v., nielsen, t.f.d. & tegner, c. 2011: silicate liquid immiscibility within the crystal mush: late-stage magmatic microstructures in the skaergaard intrusion, east greenland. journal of petrology 52, 175–222. holwell d.a., keays, r.r., mcdonald, i. & williams, m.r. 2015: extreme enrichment of se, te, pge and au in cu sulfide microdroplets: evidence from la-icp-ms analysis of sulfides in the skaergaard intrusion, east greenland. contributions to mineralogy and petrology 170:53, http://dx.doi.org/10.1007/s00410-015-1203-y irvine, t.n., andersen, j.c.ø. & brooks, c.k. 1998: included blocks (and blocks within blocks) in the skaergaard intrusion: geological relations and the origins of rhythmic modally graded layers. geological society of america bulletin 110, 1398–1447. jakobsen, j.k., veksler, i.v., tegner, c. & brooks, c.k. 2011: crystallization of the skaergaard intrusion from an emulsion of immiscible ironand silica-rich liquids: evidence from melt inclusions in plagioclase. journal of petrology 52, 345–373. keays, r.r. & tegner, c. 2015: magma chamber processes in the formation of the low-sulphide magmatic au–pge mineralization of the hot and liquid interior only tiny crystals no significant crystal settling host rocks inward-migrating mush zone subjected to in situ fractionation and crystallisation solid gabbro with crystallisation in trapped liquid fig. 4. principles of the proposed model of inward migration of mush zone and liquidus front. the inward-migrating mush zone is shown in orange. the residual bulk liquid in the centre has only small and suspended crystals of liquidus phases. the remaining bulk melt is always at liquidus due to feedback from the mush zone (see nielsen et al. 2015 for details of model). 6262 platinova reef in the skaergaard intrusion, east greenland. journal of petrology 56, 2319–2340. langmuir, c. h. 1989: geochemical consequences of in situ crystallization. nature 340, 199–205. mcbirney, a.r. 1996: the skaergaard intrusion. in: cawthorn, r.g. (ed.): layered intrusions, 147–180. amsterdam: elsevier. mcbirney,a,r. & noyes, r.m. 1979: crystallization and layering of the skaergaard intrusion. journal of petrololgy 20(3), 487–554. mckenzie, d. 2011: compaction and crystallization in magma chambers: towards a model of the skaergaard intrusion. journal of petrology 52(5), 905–930. namur, o., humphreys, m.c.s. & holness, m.b. 2014: crystallization of interstitial liquid and latent heat buffering in solidifying gabbros: skaergaard intrusion, greenland. journal of petrology 55(7), 1389–1427. nielsen, t.f.d. 2004: the shape and volume of the skaergaard intrusion, greenland: implications for mass balance and bulk composition. journal of petrology 45(3), 507–530. nielsen, t.f.d., andersen, j.c.ø. & brooks, c.k. 2005: the platinova reef of the skaergaard intrusion. in: mungall, j.e. (ed.): exploration for platinum-group element deposits, 431–455. mac short course 35. ottawa: mineralogical association of canada. nielsen t.f.d., andersen j.c.ø., holness, m.b., keiding, j.k., rudashevsky, n.s., rudashevsky, v.n., salmonsen, l.p., tegner, c. &. veksler, i.v. 2015: the skaergaard pge and gold deposit: the result of in situ fractionation, sulphide saturation, and magma chamber-scale precious metal redistribution by immiscible fe-rich melt. journal of petrology 56(8), 1643–1676. rudashevsky, n.s., rudashevsky, v.n. & nielsen, t.f.d. 2015: intermetallic compounds, copper and palladium alloys in au–pd ore of the skaergaard pluton, greenland. geology of ore deposits 57(8), 674–690. salmonsen, l.p. & tegner, c. 2013: crystallization sequence of the upper border series of the skaergaard intrusion: revised subdivision and implications for chamber-scale magma homogeneity. contributions to mineralogy and petrology 165, 1155–1171. svennevig, k. & guarnieri, p. 2012: from 3d mapping to 3d modelling: a case study from the skaergaard intrusion, southern east greenland. geological survey of denmark and greenland bulletin 26, 57–60. tegner, c., thy, p., holness, m.b., jakobsen, j.k. & lesher, c.e. 2009: differentiation and compaction in the skaergaard intrusion. journal of petrology 50(5), 813–840. wager, l.r. & brown, g.m. 1968: layered igneous rocks. edinburgh and london: oliver & boyd, 588 pp. wotzlaw, j.-f., bindeman, i.n., schaltegger, u., brooks, c.k. & naslund, h.r. 2012: high-resolution insights into episodes of crystallization, hydrothermal alteration and remelting in the skaergaard intrusive complex. earth and planetary science letters 355–356, 199–212. author’s address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark; e-mail: tfn@geus.dk geological survey of denmark and greenland bulletin 17, 2009, 25-28 the danish north sea coast is a dynamic sedimentary environment experiencing erosion, transport and re-deposition of sand along the coast. because of the natural and economic value of the coastal zone expensive protection measures such as nourishment of the coast are undertaken. the present study utilises provenance analysis techniques developed at the geological survey of denmark and greenland (geus) to characterise the coastal sand bodies by fingerprinting the heavy minerals in the sand. the aims of the study are to test these new methods in an active sedimentary environment and to develop an understanding of transport pathways along the coast. a total of c. 40 samples have been collected and analysed as part of the project. this paper gives an outline of the project and provides examples of the methods used based on six samples from the husby profile on the west coast of jylland (fig. 1). the study is a collaboration project involving geus and the department of geography and geology (dgg) at copenhagen university; geus is responsible for the analyses and dgg for sample collection. provenance analysis based on modal abundance and composition of heavy minerals is used to understand the dispersal of sand in ancient siliciclastic systems, with focus on problems relevant to the petroleum industry (e.g. larsen et al. 2006; morton et al. 2007). investigation of present-day processes and sedimentary environments may provide a key to understand how provenance indicators can be used to describe and interpret fossil clastic sedimentary systems. it is well known that properties such as grain shape and density are important factors in the processes controlling transport and deposition of mineral grains. as the heavy minerals used in provenance analysis have a wide range of densities and shapes, it is important to include these properties in studies of their dynamics in the environment. therefore geus has focused on developing computer-controlled scanning electron microscopy (cc sem), which provides information on mineralogy and mineral chemistry together with grain size and shape (keulen et al. 2008). furthermore, the development of laser ablation inductively coupled mass spectrometry (la-icp-ms) analysis has focused on a single, robust mineral species, zircon. these analyses also yield information on the age of the zircon mineral grains, and the age distribution of zircon grains is used to fingerprint the sand (knudsen et al. 2005; frei et al. 2006). saye & pye (2005) have described variations in the bulk chemical composition and particle size of coastal sands in denmark. the work outlined here is a continuation of this work, but using different methods focusing on the modal mineralogy of the heavy minerals in the sand as well as the age distribution of detrital zircon grains. sample sites samples have been collected from three settings. (1) from coastal cliffs that are being actively eroded and accordingly supply sediment to the littoral drift system along the coast of jylland. the sampling aims at fingerprinting some of the potential sources of sand mainly from till beds and glacio-fluvial sediments. (2) from four profiles orientated perpendicular to the coast; the samples represent lower shoreface, upper shoreface, beach and dunes (profiles at bovbjerg, husby, vejers and skallingen; fig. 1). the aim is to study the variation of the sediment fingerprint between the different sedimentary facies along the coast as well as the variation within identical sedimentary environments. the profiles are situated along the southward-directed net littoral drift of sand that exists between bovbjerg and skallingen. (3) from the vejers/ skallingen/grå dyb area samples will be analysed to investigate if the littoral drift sediment bypasses horns rev and to characterise the sedifingerprinting sediments along the west coast of jylland: interpreting provenance data christian knudsen, thomas kokfelt, troels aagaard, jesper bartholdy and morten pejrup © geus, 2009. geological survey of denmark and greenland bulletin 17, 25–28. available at: www.geus.dk/publications/bull 25 fig. 1. map of western denmark showing the location of the husby area (red box) and place names mentioned in the text. the map to the right shows the husby area with sample sites marked. skallingen grådyb vejers horns rev husby jylland bovbjerg 50 km vadehavet n orth sea husby 8°9’e 56°17’n 70248/70249 (black till) 70230 (dune) 70231 (swash zone) 70228 (–6 m) 70229 (–3 m) sea beach dunes other land roads ditches 1 km north sea rosa_2008:rosa-2008 01/07/09 15:48 side 25 ment exchange between the north sea and vadehavet (danish wadden sea) relative to other sources and sinks in the area. here we present data from the profile at husby. computer-controlled scanning electron microscopy (ccsem) ccsem is used at geus to analyse the composition and properties of detrital mineral grains. sediment samples are sieved and the fraction between 45 and 710 µm analysed. the samples are separated using heavy liquid (2.89 g/cm3) and the heavy mineral fraction is used for analysis. the grains are mounted in epoxy in such a way that the grains do not touch one another. the mount is polished and analysed by ccsem to determine the chemical composition of c. 1200 individual heavy mineral grains, together with their size and shape. the result of the chemical analysis of each grain is compared with a library of mineral compositions. the results are stored in a database and properties such as modal mineralogy, mineral chemistry and grain-size distributions of the individual species can be displayed (figs 2, 3). the ccsem analytical procedure is described in further detail by knudsen et al. (2005), keulen et al. (2008) and bernstein et al. (2008). la-icp-ms fingerprinting of zircon when using the modal proportions of heavy minerals for characterising the provenance of the sand, it is important to recognise potential hydrodynamic effects due to, e.g. grain density. in addition, the size and shape characteristics of the grains of the different heavy mineral species may influence their response to sedimentary processes. further more, the relative sensitivity of the heavy mineral species to 26 the 63–125 µm fraction contains c. 60% of the zircon c. 40% of the ilmenite c. 35% of the garnet c. 20% of the mafic silicates the 63–125 µm fraction contains c. 80% of the zircon c. 80% of the ilmenite c. 80% of the garnet c. 75% of the mafic silicates grain diameter (µm) 10 1006345 125 710 1000 0 50 100 50 100 c um ul at iv e w t% c um ul at iv e w t% husby –6 m husby –3 m garnet ilmenite mafic silicates zircon fig. 2. grain-size distribution curves for zircon, garnet, ilmenite and mafic silicates in two samples from husby taken at 3 and 6 m water depth. note that if only the 63 to 125 µm fraction of the heavy mineral fraction was analysed, the modal proportions of the heavy mineral species would not represent the actual proportions in the sand. fig. 3. the heavy mineral proportions in the six samples from husby recorded by ccsem. the samples from the shoreface (–3, –6 m) are characterised by low contents of garnet (c. 7%) in contrast to the samples from the beach and dune that are characterised by high garnet contents (c. 20%). 0 20 40 60 80 100 beach dune ti magnetite magnetite silicates, other mafic silicates staurolite sillimanite garnet zircon rutile leucoxene ilmenite fr ac tio n (% ) –6 m –3 m till 1 till 2 rosa_2008:rosa-2008 01/07/09 15:48 side 26 chemical and physical decomposition depends on the stability of the different minerals. to overcome these problems, it is advisable in provenance analysis to focus analysis on one mineral species because differences in hydrodynamic behaviour or weathering can thereby be discounted. for this purpose the extremely stable mineral zircon is very well suited. a key property of zircon in provenance analysis is the relatively high content of uranium, which makes the mineral well suited for dating. at geus the measurements are carried out on a laser ablation inductively coupled mass spectrometer (la-icp-ms; frei et al. 2006; frei & gerdes 2009). the sand is physically separated using a shaking table to concentrate the very heavy minerals. from this concentrate, the zircon is picked out and mounted (c. 150 zircon grains per sample) in epoxy. the mount is polished and introduced to the laser ablation system. a c. 30 µm diameter spot is ablated with the laser in the core of each zircon grain, and the ablated material is introduced into the icp-ms, in which the ratios between the different u and pb isotopes are measured. results from husby some results of analyses of six samples collected near husby (fig. 1) are discussed here. two samples are from till beds exposed in the coastal cliff, one is from a dune, one is from the active swash zone and two samples are from the shoreface, at water depths of 3 and 6 m. the grain-size distribution varies among the different heavy mineral species (fig. 2). zircon with the highest density is finer-grained than lighter minerals such as ilmenite and garnet, and the mafic silicates (amphibole, pyroxene and epidote) are the most coarse-grained of the heavy minerals. this is in close accordance with that predicted by stokes’ law, and implies that grain-size sorting has acted according to grainfall velocity. the steepness of the grain-size distribution curves also varies among the different minerals. the garnet curve is steeper and narrower than that of, for example, 27 0 500 1000 1500 2000 2500 3000 3500 4000 age (ma) 0 5 10 15 20 25 30 fr eq ue nc y 70249 husby (till 2) n = 91/113, 90–110% conc. r el at iv e pr ob ab ili ty 0 5 10 15 20 fr eq ue nc y 70248 husby (till 1) n = 108/139, 90–110% conc. r el at iv e pr ob ab ili ty 0 5 10 15 20 fr eq ue nc y 70230 husby (dune) n = 92/92, 90–110% conc. r el at iv e pr ob ab ili ty 0 5 10 15 20 25 30 35 fr eq ue nc y 70231 husby (swash zone) n = 123/124, 90–110% conc. r el at iv e pr ob ab ili ty 0 10 20 30 40 fr eq ue nc y 70229 husby (–3 m) n = 146/147, 90–110% conc. r el at iv e pr ob ab ili ty 0 5 10 15 20 25 fr eq ue nc y 70228 husby (–6 m) n = 129/129, 90–110% conc. r el at iv e pr ob ab ili ty fig. 4. age distribution of zircons in the samples from husby. the frequency (red histograms) indicates the number of zircons in the age brackets (100 ma). the green area indicates the content of concordant grains whereas the yellow area indicates the discordant grains (only present in the till samples). the samples from the shoreface (–3, –6 m) are characterised by two main populations, one with ages between 1000 ma and 1300 ma and one with ages between 1400 ma and 1900 ma, the latter being the most abundant. samples from the beach and dune contain the same populations, but are characterised by more equal proportions of the two populations. it can also be noticed that the samples from the till in the cliff behind the beach are characterised by a pattern similar to the samples from the shoreface. rosa_2008:rosa-2008 01/07/09 15:48 side 27 ilmenite or mafic silicates. this could be caused by differences in grain shape, and one of the aims of the project is to investigate how the grain-size distribution varies among the different sedimentary environments along the west coast of jylland. in provenance analysis it is often assumed that ratios between the heavy mineral species can be used as provenance indicators. due to analytical constraints, or chosen standard procedures, a given size-fraction (typically 63 to 125 µm; hallsworth & chisholm 2008; yang et al. 2009) is often used for such analysis, regardless of the fact that the sands which are compared have different overall grain-size distributions. as can be seen on fig. 2, for example, the modal proportions of the heavy minerals would not reflect the true values if only the fraction between 63 and 125 µm was analysed. the sample husby –6 m would be characterised by its finer part, with zircon being over-represented whereas the sample husby –3 m would be close to the modal value. in fig. 3, the modal proportions of the heavy minerals show that the composition of the heavy mineral fraction in the shoreface sands (–6 and –3 m) is similar but different to the beach sand and dune sands, which in turn resemble each other. the main difference between these two pairs is that garnet is more abundant in the beach and dune sands than in the shoreface sand. the heavy mineral assemblages from the tills exposed in the cliff are different to the beach and dune sands, but similar to the sand from the shoreface. figure 4 shows the age distributions of the zircon grains in the six samples. the samples show roughly the same spectra: a few archaean (>2500 ma) grains, an early proterozoic maximum (c. 2000–1500 ma), a middle proterozoic maximum (c. 1400–900 ma) and a caledonian peak (c. 400–300 ma). it is a characteristic feature of the shoreface sands and the till samples that early proterozoic zircons are more abundant than middle proterozoic zircons. in contrast, the two populations are equally represented in the beach and dune samples. the middle proterozoic zircons were probably de rived from the sveco-norwegian orogen in southern norway and western sweden to the north of jylland whereas the early proterozoic zircons were probably derived from jotnian and sveco-fennian sources to the east. both the mineral paragenesis data and the zircon age distributions suggest that the shoreface sands have a provenance that is distinct from the beach and dune sands at husby. this is surprising, and further work will concentrate on explaining these patterns, with the aim of understanding the processes responsible for the observed distribution of these sands. in addition, the distribution of heavy mineral assemblages along the shore will be analysed. references bernstein, s., frei, d., mclimans, r.k., knudsen, c. & vasudev, v.n. 2008: application of ccsem to heavy mineral deposits: source of high-ti ilmenite sand deposits of south kerala beaches, sw india. journal of geochemical exploration 96, 25–42. frei, d., hollis, j.a., gerdes, a., harlov, d., karlsson, c., vasquez, p., franz, g., johansson, l. & knudsen, c. 2006: advanced in situ geo chronological and trace element microanalysis by laser ablation techniques. geological survey of denmark and greenland bulletin 10, 25–28. frei, d. & gerdes, a. 2008: precise and accurate in situ u-pb dating of zircon with high sample throughput by automated la-sf-icp-ms. chemical geology 261, 261–270. hallsworth, c.r. & chisholm j.i. 2008: provenance of late carboniferous sandstones in the pennine basin (uk) from combined heavy mineral, garnet geochemistry and palaeocurrent studies. sedimentary geology 203, 196–221. keulen, n.t., frei, d., bernstein, s., hutchison, m.t., knudsen, c. & jensen, l. 2008: fully automated analysis of grain chemistry, size and morphology by ccsem: examples from cement production and diamond exploration. geological survey of denmark and greenland bulletin 15, 93–96. knudsen, c., frei, d., rasmussen, t., rasmussen, e. s. & mclimans, r. 2005: new methods in provenance studies based on heavy minerals: an example from miocene sands in jylland, denmark. geological survey of denmark and greenland bulletin 7, 29–32. larsen, m., knudsen, c., frei, d, frei, m., rasmussen, t. & whitham, a.g. 2006: east greenland and faroe–shetland sediment provenance and palaeogene sand dispersal systems. geological survey of denmark and greenland bulletin 10, 29–32. morton, a.c., herries, r. & fanning, c.m. 2007: correlation of triassic sandstones in the strathmore field, west of shetland, using heavy mineral provenance signatures. in: mange, m. & wright, d.t. (eds): heavy minerals in use. developments in sedimentology 58, 1037–1072. saye, s.e. & pye, k. 2006: variations in chemical composition and particle size of dune sediments along the west coast of jutland, denmark. sedimentary geology 183, 217–242. yang, s., wang, z., guo, y., li, c. & cai, j. 2009: heavy mineral compositions of the changjiang (yangtze river) sediments and their provenance-tracing implication. journal of asian earth sciences 35, 56–65. authors’ addresses c.k. & t.f.k., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ckn@geus.dk t.a., j.b. & m.p., department of geography and geology, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. 28 rosa_2008:rosa-2008 01/07/09 15:48 side 28 geological survey of denmark and greenland bulletin 35, 2016, 23-26 23 the danish subsurface contains deep geothermal resources which may contribute for hundreds of years to the mixed danish energy supply (mathiesen et al. 2009). at present only a limited fraction of these resources are utilised in three existing geothermal power plants in thisted, margretheholm and sønderborg (fig. 1) where warm formation water is pumped to the surface from a production well and, after heat extraction, returned to the subsurface in injection wells (fig. 2). deep geothermal energy has the advantage of being a sustainable and environmentally friendly energy source which is furthermore independent of climate and seasonal variations, in contrast to wind and solar energy. the implementation of deep geothermal energy for district heating replacing conventional energy sources, especially coal and oil, may thus lead to a considerable reduction in the emission of greenhouse gases. there are therefore good reasons to include geothermal energy as a central component in denmark’s future supply of energy for district heating. furthermore, heat-demanding industries may consider the possibility to integrate geothermal energy and energy storage in their production process. in order to facilitate the use of geothermal energy, a broad majority in the danish parliament has granted financial support for initiatives within the geothermal field (energy policy agreement of march 22, 2012). the present paper deals with one of the outcomes of this agreement, namely a webgis portal with an overview of existing and a webgis portal for exploration of deep geothermal energy based on geological and geophysical data henrik vosgerau, anders mathiesen, morten sparre andersen, lars ole boldreel, morten leth hjuler, elina kamla, lars kristensen, christian brogaard pedersen, bjarni pjetursson and lars henrik nielsen margretheholm sønderborg thisted margretheholm sønderborg thisted geothermal potential no potential one reservoir two or more reservoirs structurally complex area and limited seismic data fig. 1. distribution of lithostratigraphic units with reservoir properties suitable for geothermal exploration in the geothermal depth zone (800–3000 m), see the main text. the red line indicates the transect covered by the geosection shown in fig. 5. positions of the operating geothermal plants are shown. fig. 2. principle sketch of a geothermal power plant with production and injection wells to a deep-seated geothermal sandstone reservoir. © 2016 geus. geological survey of denmark and greenland bulletin 35, 23–26. open access: www.geus.dk/publications/bull te m p er at u re g ra d ie n t 2 5 – 3 0 °c /k m sasasasasasasasasaasasasas ndndndndndndndndndndnd aaaaaaaaandndndndndndndndndddnnd gggggggggggggrararararararrararararavevevevevevevevveveveev llllllll clclclclclclclccclayayayayayayayayaya tttttttttttttilililililililili llllllllllll chchchchchchchchchchchchchchchhhalalalalalalalalaalallalla kkkkkkkkkkkkkkkk clclclclclclclclclclclccclclclcllclayayayayayayayayaayaaayayayyaystststststststststststttsttttononononononononononononononooo eeeeeeeeeeeeeee sasasasasasasasasasaasandndndndndndndndndddndn stststststststststttononononononononnonnneeeeeeeeee gegegegegegeggegegegeeeototototototototooo hehehehehehehheeheermrmrmrmrmrmmrmrmmmr alalalalalalalal rererererereerrerrrrr seseseseseseeseseeseservrvrvrvrvrvrvrvrvrvoioioioioioiooioiooo rrrrrrrrrr ~1,5 km 2,5 km ~75°c p ro d u ct io n w el l in je ct io n w e ll 2424 interpreted geological and geophysical data. this will be relevant for all stakeholders in the exploration of deep geothermal resources in the danish subsurface. the portal focuses on geothermal reservoirs within the 800–3000 m depth interval and provides an overview of the amount and quality of existing geodata, the geological composition of the subsurface, and interpreted thematic products such as geological maps of potential geothermal reservoirs. a comprehensive map from the portal showing onshore and nearoffshore locations where the geological conditions are potentially suitable for extraction of deep geothermal energy in denmark is shown in fig. 1. many of the thematic maps are outcomes of the project the geothermal energy potential in denmark – reservoir properties, temperature distribution and models for utilization under the programme sustainable energy and environment funded by the danish agency for science, technology and innovation. geological requirements for geothermal exploitation in denmark, successful geothermal exploitation in the deep subsurface requires the presence of thick and laterally coherent sandstone reservoirs with high porosity and permeability, which can ensure effective extraction and re-injection of formation water. a thick and coherent reservoir which is not hydraulically compartmentalised by faults, lateral lithological changes (e.g. grain size) and/or diagenetic features such as compaction and mineralisation implies that a large volume of warm water may be accessible, and that production and injection wells can be placed at appropriate distances from each other while remaining hydraulically connected. a certain distance, e.g. 1,5 km, is needed between the filter screens at reservoir level in order to delay the arrival time of the cool, re-injected water to the vicinity of the production well (fig. 2). as a rule of thumb this delay should be more than 30 years. on its way to the production well, the cooled water is reheated to some extent by the reservoir matrix and by heat transfer from poorly or non-permeable boundary strata (usually claystone) above and below the reservoir sandstones. another important requirement is to find areas where the product of temperature and extractable water volume is large enough to ensure an economically viable geothermal plant. the temperature–depth gradient of 25–30°c/km in the danish subsurface implies that at depths shallower seismic data coverage and quality very poor poor resonable good very good petrophysical log data quality no data or poor uncertain good or resonable seismic data coverage and quality very poor poor reasonable good very good petrophysical log data and quality data absent or poor uncertain reasonable or good fig. 3. coverage and quality of seismic and petrophysical log data from deep wells. the quality indexes reflect to which degree the data can be used to extract information about geothermal relevant lithostratigraphic units in the deep subsurface. 25 than 800 m the temperature is generally too low, whereas at depths greater than 3000 m, diagenetic alterations related to high pressure–temperature conditions reduce the porosity and permeability of the reservoir sandstones (poulsen et al. 2015; kristensen et al. 2016). hence, the focus of the portal is the 800–3000 m depth interval. the geological database the geological data that provide information about denmark’s deep subsurface mainly consist of information from deep wells and seismic surveys collected over a number of years during oil and gas exploration, and to a lesser extent during studies of potential gas storage and geothermal exploration. the geographical coverage and quality of the data vary considerably as outlined in fig. 3. generally, it is possible to obtain detailed information about the penetrated sedimentary successions from the well data including the depth, thickness and reservoir properties of sandstone reservoirs, e.g. based on core data and petrophysical evaluation of well log data (fig. 4). the seismic data have been used for largescale mapping of the depth, thickness and lateral extent of lithostratigraphic units known to contain geothermal reservoir sandstones, as well as for identification and mapping of major faults (fig. 5). this work involved compilation and integration of the many seismic surveys of different age and quality into a coherent seismic network. then the network was interpreted in its entirety and used to generate nationwide depth maps to important subsurface horizons, with conversions from seismic-wave travel time to depths below sea level. g as su m f o rm at io n -1580 -1480 -1600 -1500 -1520 -1540 -1560 -1460 1580 1600 1500 1620 1520 1640 1540 1660 1560 gr api0 250 6 160 ms/ft dt 60 inches cali 16 10000 md perm_log 1 prs nphi 0.6 v/v dec.0 rhob 1.7 3g/cm 2.7 phie 0 % 40 metres m (tvdss) coresdepth (m md) raw logs permeability porosity sandstone shale coal potential reservoir sand fig. 4. petrophysical log evaluation of the gassum formation in the stenlille-1 well including interpretations of lithology, porosity and permeability. gr: gamma-ray log. api: gamma ray radioactivity expressed in accordance with the american petroleum institute. cali: caliper log. dt: sonic log. prs: potential reservoir sand. nphi: total porosity log. rhob: density log. perm_log: permeability log. phie: effective porosity log. md: measured depth. tvdss: vertical depth below sea level. danian limestone and chalk group haldager sand fm frederikshavn fm fjerritslev fm gassum fm incl. lower jurassic unit fault top pre-zechstein skagerrak fm (skagerrak–kattegat platform) bunter sandstone fm and skagerrak fm (time equivalent subinterval to bunter sandstone fm) zechstein group skagerrak– kattegat platform sorgenfrei–tornquist zonedanish basinringkøbing–fyn high 0 -1 -2 -3 -4 t w t ( s) nesw nøvling-1 nøvling-1 flyvbjerg-1 flyvbjerg-1 frederikshavn-1 frederikshavn-1 haldager-1 haldager-1 sæby-1 sæby-1 25 km fig. 5. a regional sw–ne geosection through central and northern jylland as outlined in fig. 1. the profile is constructed on the basis of composite seismic profiles shown in the background. mapped lithostratigraphic units with individual treatment in the portal are highlighted. these include those units which may contain geothermal sandstone reservoirs. in several places the units are truncated by faults, and vertical salt movement from the zechstein group has in places uplifted or penetrated the overlying succession. major structural elements are indicated. depth is given in seismic two-way travel time (s). 2626 the depth maps were subsequently used as a base for a new set of maps where constrained reservoir properties such as porosity and temperature were added in relation to depth. in general, the amount of data is too limited for traditional statistical uncertainty analysis. however, based on quality and spatial density analysis and geus’ general experience, rough uncertainty intervals for the estimated reservoir values are presented for each well in the portal. the webgis portal – content, functionality and perspectives the first map in the webgis portal shows those areas where lithostratigraphic units, containing sandstones with sufficient reservoir properties are considered to be present within the depth interval of 800–3000 m (fig. 1). the lithostratigraphic units are also treated separately and linked with several relevant theme maps showing e.g. depth, thickness, major faults, reservoir transmissivity and temperature. additional data, available in the portal, can be added to each of these maps, such as the distribution and quality of seismic data and petrophysical log data from deep wells. the depth maps can also be visualised through an interactive 3d viewer providing an intuitive overview of the subsurface topography of a selection of important horizons (fig. 6). these maps are to be considered as indicative and are only meant for regional use. this is because the geological data in many areas only provide a rough picture of the subsurface, especially where data are sparse and/or of poor quality. in addition, some of the maps are based on several stages of approximations and generalisations, for instance where largely depth-depending reservoir parameters are paired with depth maps derived from seismic surveys, as briefly described above. the maps are not final. new well and seismic data can be added and new interpretation tools such as seismic interpolation tools can be included. refined geological models may also lead to modifications. however, the present maps give a good indication of where in denmark deep geothermal exploration is relevant, as seen from the geological prerequisites. the use of the maps may thus ensure that new geothermal exploration is directed towards those areas that appear to be most promising based on current knowledge. the various geological map themes also form an important basis for an initial analysis of the geothermal potential at a specific site where the construction of a geothermal plant is considered. in this first step, the maps give an overview of the potential geothermal reservoir intervals (lithostratigraphic units) that may be relevant at the site, and information regarding the type, amount and quality of existing geological data. a more comprehensive estimate of the geothermal potential in a specific area must be based on detailed analysis of the local data and incorporation into local geological models. this work model has been used in an assessment of 28 potential geothermal sites in denmark in the so-called screening project; another outcome of the above-mentioned policy agreement. the results of this work are also accessible free of charge from the webgis portal at http://dybgeotermi.geus.dk. references kristensen, l., hjuler, m.l., frykman, p., olivarius, m., weibel, r., nielsen, l.h. & mathiesen, a. 2016: pre-drilling assessments of average porosity and permeability in the geothermal reservoirs of the danish area. geotherm energy 4:6, 2–27, http://dx.doi.org/ 10.1186/ s40517-016-0048-6 mathiesen, a., kristensen, l., bidstrup, t. & nielsen, l.h. 2009: vurdering af det geotermiske potentiale i danmark. danmarks og grønlands geologiske undersøgelse rapport 2009/59, 30 pp. poulsen, s.e., balling, n. & nielsen, s.b. 2015: a parametric study of the thermal recharge of low enthalpy geothermal reservoirs. geothermics 53, 464–478. fig. 6. interactive 3d tool, available in the webgis portal, visualising the variation of subsurface topography of mapped seismic horizons in a given 10 × 10 km area around a selected point. authors’ address geological survey of denmark and greenland, øster voldgade 20, dk-1350 copenhagen k, denmark,. e-mail: hv@geus.dk geological survey of denmark and greenland bulletin 11, 9-31 9 evolution of neoarchaean supracrustal belts at the northern margin of the north atlantic craton, west greenland julie a. hollis, marie keiding, bo møller stensgaard, jeroen a.m. van gool and adam a. garde the archaean north atlantic craton of west greenland collided at c. 1.9 ga with a lesser-known archaean craton to the north, to form the nagssugtoqidian orogen. the palaeoproterozoic metamorphic grade and strain intensity decrease northward through the orogen, allowing investigation of the reworked archaean components in its northern part. two archaean supracrustal belts in this region – the ikamiut and kangilinaaq belts – are investigated here using field mapping, aeromagnetic data, zircon geochronology, and geochemistry. both belts comprise quartzo-feldspathic and pelitic metasedimentary rocks, amphibolite, and minor calc-silicate rocks, anorthosite and ultramafic rocks. pbpb and u-pb dating of detrital zircons and host orthogneisses suggest deposition at c. 2800 ma (kangilinaaq belt) and after 2740 ma (ikamiut belt); both belts have zircons with neoarchaean metamorphic rims. metasedimentary rocks and orthogneisses at ikamiut share similar steep ree signatures with strong lree enrichment, consistent with local derivation of the sediment and deposition directly onto or proximal to the regional orthogneiss precursors. zircon age data from kangilinaaq indicate both local and distal sources for the sediment there. geochemical data for kangilinaaq amphibolites indicate bimodal, mixed felsic–mafic source rocks with island-arc basaltic affinities, consistent with a shelf or arc setting. both belts experienced a similar tectono-metamorphic history involving neoarchaean amphibolite facies peak metamorphism at c. 2740–2700 ma, possibly due to continued emplacement of tonalitic and granodioritic magmas. nagssugtoqidian lower amphibolite facies metamorphism at c. 1850 ma was associated with development of the large-scale f 2 folds and shear zones that control the present outcrop pattern. the observed differences in the sources of the kangilinaaq and ikamiut belts and their shared post-archaean history suggest they were formed in different neoarchaean environments proximal to and on a continental plate, and were amalgamated in a convergent margin setting shortly after their deposition. keywords: north atlantic craton, northern nagssugtoqidian orogen, la-icp-ms, sims, zircon ________________________________________________________________________________________________________________________________________________________________ j.a.h., b.m.s., j.a.m.v.g. & a.a.g., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: jho@geus.dk m.k., geological museum, university of copenhagen, øster voldgade 5–7, dk-1350 copenhagen k, denmark. © geus, 2006. geological survey of denmark and greenland bulletin 11, 9–31. available at: www.geus.dk/publications/bull 10 greenland inland ice greenland canada archaean, variably reworked metasedimentary rocks surficial deposits basalt quaternary cretaceous–palaeogene proterozoic sandstone sarfartoq carbonatite complex sisimiut charnockite arfersiorfik quartz diorite granodioritic and granitic gneiss granodioritic and granitic gneiss orthogneiss dioritic gneiss orthogneiss metasedimentary rocks amphibolite (including proterozoic components) intermediate to basic intrusions boye sø anorthosite complex archaean, unreworked thrustt t jakobshavn isfjord ussu itnordre strømfjord arfersiorfik aasiaat qeqertarsuaq ilulissat qasigiannguit kangaatsiaq attu sisimiut sø nd re str øm fjord kangerlussuaq naternaq nuussuaq vaigat 51° disko n ag ss ug to qi di an o ro ge n r in ki an fo ld b el t n or th a tl an ti c cr at on sn o c n o nssb itz n n o inland ice fig. 2a ikamiut region fig. 2b kangilinaaq peninsula sydostbugten 0 50 km tt t t t t t t t t t t ttt t disko bugt sukkertoppen iskappe 68° 66° 70° fig. 1. geological map of the nagssugtoqidian orogen, west greenland, from van gool et al. (2002b). frames show locations of figs 2 and 3. 11 the palaeoproterozoic nagssugtoqidian orogen, central west greenland, comprises archaean and less abundant proterozoic orthogneiss and metasedimentary rocks deformed and metamorphosed at c. 1850 ma during collision of the north atlantic craton with a lesser-known, likewise archaean craton to the north (kalsbeek et al. 1987; taylor & kalsbeek 1990; kalsbeek & nutman 1996; connelly et al. 2000; van gool et al. 2002a). the orogen extends from søndre strømfjord in the south, northward to disko bugt and possibly farther into the largely contemporaneous rinkian fold belt (fig. 1). the metamorphic grade associated with orogenesis decreases from granulite facies in the collisional core (the central nagssugtoqidian orogen) to amphibolite facies in the southern foreland and the northern part of the orogen. also the penetrative palaeoproterozoic deformation diminishes toward the north in the northern nagssugtoqidian orogen (nno), and heterogeneous strain distribution may have been important in the preservation of pre-nagssugtoqidian, i.e. archaean structural fabrics and metamorphic assemblages and textures (van gool et al. 2002a; garde et al. 2004; hollis et al. 2004; piazolo et al. 2004; mazur et al. 2006, this volume). as a consequence of the northward decrease in the palaeoproterozoic thermal overprint and deformation, the nno provides an opportunity for investigation of the pre-nagssugtoqidian history of its archaean components. in particular, its supracrustal belts can provide valuable information on the tectonic environment(s) of their formation, their relationship to the plate-tectonic configuration, and whether different parts of the craton experienced the same or different archaean histories. here we investigate two supracrustal belts within the nno – the kilometre-wide ikamiut and kangilinaaq belts – that crop out on the western and eastern sides of sydostbugten in southern disko bugt (fig. 1). parts of the nno were mapped by the geological survey of greenland in the 1960s for its 1:500 000 scale geological map series (noe-nygaard & ramberg 1961; henderson 1969) and also by the geological survey of denmark and greenland (geus) in 2000–2003 for the 1:100 000 scale geological map series (see below). henderson (1969) identified a complex map-scale fold structure that dominates the ikamiut peninsula and adjacent inland areas in the western sydostbugten region, and also outlined many of the dominant lithologies and large structural elements in the kangilinaaq region. in this paper we present geological, geochemical, geochronological and geophysical data from work carried out in the period 2000–2003 by geus mapping teams for the kangersuneq and ikamiut 1:100 000 scale geological map sheets (van gool 2005; garde in press); part of this work is reported in more detail in keiding (2004). aeromagnetic data for the sydostbugten region are correlated with major lithological and structural elements. geochemical data from amphibolites in the kangilinaaq region interpreted as deformed and metamorphosed basaltic volcanic rocks, and interlayered pelitic rocks, are used to determine the likely depositional environment. zircon pb-pb and u-pb geochronology on granodioritic orthogneisses and a metasedimentary rock from the ikamiut region is compared with existing data from the ikamiut and kangilinaaq regions. finally, the implications for regional neoarchaean tectonics are discussed. ikamiut belt and host rocks west of sydostbugten the ikamiut belt is a deformed, kilometre-thick sequence of biotite schists, with less abundant siliceous and pelitic rocks, amphibolite and minor ultramafic rocks. the belt forms a ten kilometre-scale antiform in the north-western part of sydostbugten (fig. 2a). it is everywhere in contact with c. 2830–2760 ma old, tonalitic to granodioritic orthogneiss (pb-pb whole rock, kalsbeek et al. 1987; u-pb zircon, connelly & mengel 2000 and this study), which dominates the region. the original nature of the contacts between the supracrustal belt and the regional orthogneiss is obscured by later ductile deformation (see also østergaard et al. 2002). rb-sr data for 12 metasedimentary samples from this belt, near ikamiut, gave an age of c. 1880 ma for closure of the rb-sr system and a very high initial 87sr-86sr ratio of c. 0.712, suggesting that these rocks were deposited at around 2.8 ga and isotopically strongly reset during nagssugtoqidian metamorphism (kalsbeek & taylor 1999). structure the structural pattern is dominated by kilometre-scale, closed, upright f 2 folds folding an s 1 foliation and associated with a moderate to intense, ene-striking s 2 foliation (fig. 2a). preservation of s 1 fabrics is found in areas of low d 2 strain, typically within the cores of large f 2 folds. outcrop-scale, parasitic f 2 folds associated with weak to moderately developed mineral lineations (l 2 ) plunge at shallow angles to the wsw. in the eastern ikamiut region, l 2 lineations are shallow and, in some cases, plunge to the east. some f 2 folds may be doubly plunging and/or refolded, consistent with localised outcrop-scale refolded 12 b 467433 467436 467401 467403 467404 467405 467408 467444 467445 467446 467413 467426 467440 467420 467423 qasigiannguit 30 10 44 35 40 78 10 14 20 20 8 6 10 10 22 25 30 35 32 38 44 48 35 44 63 63 50 52 40 35 73 68 45 13 60 60 15 55 4515 28 55 20 50 70 14 62 48 44 38 41 28 30 2631 1035 3513 39 22 2886 32 68°45' 68°55' 51°15' 51° 50°45' la ks eb ug t kan ger lulu k f1 f2 f2 f2 kan gil ina aq 68°35' 68°40' a467526 440938 44 15 40 2 3751 9 62 55 10 48 75 14 16 38 21 22 20 2 79 5 72 87 85 3212 8 34 23 8 8 14 5 5 35 2 4 22 5 nivaap paa kan ger sun eq 440910 ikamiut 52° 51°45'langesund f1 f2 f2 f2 f1 amphibolite, unspecified quartzo-feldspathic paragneiss, locally garnet bearing pelitic gneiss, garnet-biotitesillimanite bearing biotite schist, dark, well foliated, with small garnets marble and calc-silicate rock quaternary cover anorthosite mafic dyke trace of axial surface, antiformal/synformal fold 5 km homogeneous amphibolite – mafic intrusive rock complex layered amphibolite – mafic supracrustal rocks two-mica granite pegmatite porphyritic granite porphyritic orthogneiss to foliated granite porphyritic quartz diorite n 5 km n biotite-bearing, grey orthogneiss fig. 2. preliminary geological interpretation maps on (a) the ikamiut region and (b) the kangilinaaq peninsula, showing representative structural data, sample numbers and localities. for regional location see frames in fig. 1. 13 folds. the kilometre-scale, upright f 2 folds are probably parasitic on the major antiformal structure that dominates the outcrop pattern. ten to hundred metre-scale, shallowly w-plunging f 2 folds of flat-lying s 1 foliation are abundant in the tonalitic to granodioritic orthogneiss inland at the head of nivaap paa (fig. 2a). these folds are difficult to trace for long distances along strike. the inland outcrop is relatively poor, and available outcrop suggests that some of the folds die out towards the west, apparently because of homogeneity and lack of ductility contrast within the orthogneiss. lithologies, mineral assemblages and fabrics the tonalitic to granodioritic orthogneiss is a compositionally layered, medium-grained, pale pink and grey rock dominantly comprising plagioclase and quartz, with lesser k-feldspar and disseminated biotite. a mediumto coarse-grained gneissic layering (s 1 ) is discontinuous on a scale of metres to tens of metres and commonly displays intrafolial isoclinal folds. the orthogneiss typically also holds a moderate, ene-striking s 2 foliation that is a partially to completely transposed s 1 fabric, and a weakly to moderately developed, shallow w-plunging l 2 mineral lineation. mediumto coarse-grained, centimetre-scale granitic veins indicate variable d 2 strain: they form a layering that is typically transposed into s 2 , but in some cases they are slightly discordant. the supracrustal sequence is dominated by biotite schists to gneissic rocks that typically comprise plagioclase, quartz, biotite, and rarely garnet. interlayered with these rocks occur local, more mica-rich and aluminous layers up to 20 m thick, which commonly display a schistose s 1 fabric. these rocks typically comprise biotite, plagioclase, quartz, muscovite, sillimanite, and garnet, with accessory tourmaline. locally, in the most micaceous parts, a d 2 crenulation of s 1 biotite, plagioclase and quartz ± sillimanite is developed, with axial planes parallel to the dominant regional s 2 gneissosity of the tonalitic–granodioritic orthogneiss host, and fine-grained biotite and muscovite along the crenulations. aggregates of fine-grained sillimanite and biotite in biotite-rich schists form blocky, centimetre-scale patches interpreted as pseudomorphs after andalusite. this suggests that the s 1 fabrics and assemblages were formed at low-pressure (< 3.85 kbar), lower amphibolite facies metamorphic conditions, followed by increasing temperature (and possibly also pressure) into sillimanite-grade conditions. a distinct unit of siliceous paragneiss, locally garnetbearing, is also volumetrically important. it is distinguished from the biotite schist/gneiss by its more quartz-rich and mica-poor composition. it is often difficult to distinguish this lithology from the tonalitic to granodioritic orthogneiss, particularly in inland areas where outcrop is poor. amphibolite layers, which are 10–50 m thick and laterally discontinuous on a kilometre-scale, are associated with the metasedimentary sequence. they commonly occur at or near boundaries between the metasedimentary rocks and orthogneiss. the amphibolites are medium grained and comprise hornblende with lesser plagioclase and quartz, and locally clinopyroxene ± garnet. in some cases they show distinct mafic–felsic layering, and they commonly contain thin (0.5–5 cm), discontinuous felsic layers. a few isolated occurrences of intensely deformed anorthosite occur at tectonised boundaries between the tonalitic to granodioritic orthogneiss and metasedimentary rocks. the largest occurrence is in the northern island group of nivaap paa (fig. 2a). the anorthositic rock is coarse grained and ‘zebra-striped’, and consists of hornblende and calcic plagioclase with a variably developed foliation and an intense linear fabric. the mafic parts are commonly boudinaged within the more felsic component. an extensive body of mediumto coarse-grained granite is located within the hinge zone of the large antiformal structure along the southern coast of nivaap paa. the granite is porphyritic and white to red in colour, and holds a weak gneissose fabric. its northern contact with the regional orthogneiss is tectonised and possibly tectonically repeated. the relatively undeformed nature of the granite suggests it intruded into the orthogneiss after formation of the regional gneissose fabric (s 2 ). the granite contains thin lenses and layers of medium-grained amphibolite, and is bounded to the south by a layer of amphibolite 50–400 m thick, the outcrop of which defines a tight synformal fold closure (fig. 2a). kangilinaaq belt and host orthogneiss east of sydostbugten the kangilinaaq peninsula (fig. 2b) is dominated by a kilometre-scale synformal structure comprising a series of ne-trending, upright, isoclinal f 1 and f 2 folds that repeat a thick supracrustal sequence. the most common lithologies are quartzo-feldspathic and pelitic metasedimentary rocks, with lesser amphibolite and subordinate marble and calc-silicate rocks. an equivalent supracrustal sequence is found south of kangersuneq fjord, on the southern limb of an antiformal fold running through the fjord. a lithologically distinct unit of amphibolite and associated metasedimentary rocks runs through the town 14 of qasigiannguit in the west of the peninsula. for ease of reference this unit is named the qasigiannguit amphibolite in the following. it is separated from the main supracrustal sequence by 200–500 m of high-strain archaean orthogneiss. the significance of this high-strain zone in terms of the original supracrustal stratigraphy is uncertain, and thus the qasigiannguit amphibolite may or may not be part of an originally continuous supracrustal series on the kangilinaaq peninsula. for descriptive reasons the two supracrustal sequences are collectively termed the kangilinaaq belt in the following. pelitic rocks from the main supracrustal sequence contain archaean detrital zircon populations in the range 2820–2760 ma, with a minimum depositional age constrained by an intrusive two-mica granite at 2723 ± 15 ma (thrane & connelly 2006, this volume). metamorphic zircon growth occurred at 1920–1820 ma in various rocks (keiding 2004; thrane & connelly 2006, this volume). age data are addressed in more detail in the discussion. structure the structural pattern is dominated by large, tight to isoclinal folds. especially along the south-eastern side of the peninsula, the quartzo-feldspathic and pelitic rocks are intensely folded into upright folds on scales from decimetres to tens of metres with shallowly ne-plunging fold axes. older, isoclinal, often intrafolial folds indicate that the upright folds are at least second-generation (f 2 ). a shear zone containing mylonitic orthogneiss bounds the supracrustal rocks to the south-east. it can be traced from the south-western part of the peninsula to half-way up kangersuneq fjord, where it meets the water (fig. 2b). kinematic indicators and a lineation suggesting sinistral/ top to the west movement are poorly developed in the shear zone. the continuation of the shear zone may be found in a poorly developed, but continuous se-trending shear zone south of kangersuneq fjord, marked by a sliver of metasedimentary rocks. north-west of the supracrustal sequence, no similar shear zone was found, although some smaller zones of high strain were recognised. lithologies, mineral assemblages and fabrics the predominant regional lithologies are layered, grey tonalitic to granodioritic orthogneisses interleaved with supracrustal rocks. the orthogneisses contain variable proportions of plagioclase, quartz and biotite, with minor kfeldspar and hornblende. compositional layering of orthogneiss with thin amphibolite layers interpreted as highly attenuated enclaves, give the rocks a layered appearance. the orthogneisses are intersected by concordant to slightly discordant, mediumto coarse-grained, centimetreto half metre-scale granitic veins interpreted as derived from local melts. the orthogneisses show intrusive contacts into part of the supracrustal sequence (see below), although it is not certain that this relationship applies to all of the supracrustal rocks on the peninsula. the gneissic fabric of orthogneiss in the core of the peninsula is locally disturbed by pods and sweats of partial melt, which can contain millimetre-sized garnets. garnet formation in the orthogneiss is restricted to the core of the peninsula, a region of abundant metasedimentary rocks. the garnet formation may be the product of contamination during partial melting of the metasedimentary rocks, either during intrusion of the precursors to the orthogneiss, or during metamorphism. variably deformed quartz diorite occurs in two localities. typically it has tectonised contacts with the supracrustal rocks, but east of qasigiannguit it has intrusive contacts to the latter and has yielded a u-pb zircon emplacement age of 2801 ± 34 ma (thrane & connelly 2006, this volume). the main supracrustal sequence is dominated by medium-grained, quartzo-feldspathic metasedimentary rocks that commonly contain garnet. where garnet is absent in these rocks, they are difficult to distinguish from the orthogneisses. the quartzo-feldspathic paragneiss alternates on metreto 100 metre-scale with pelitic rocks, amphibolite and rare calc-silicate rocks. the pelitic rock comprises quartz, plagioclase, biotite, garnet, and sillimanite. locally, it has a large component of granitic partial melt, commonly occurring in boudinaged lenses, indicative of upper amphibolite facies metamorphic conditions. amphibolites are commonly dark and subtly layered, fine to medium grained and few metres to 50 m wide. remnants of deformed pillows are locally present. in the eastern part of the sequence isolated lenses of metamorphosed ultramafic rocks occur in a few locations, commonly within amphibolites. they comprise predominantly amphibole and orthopyroxene, with or without clinopyroxene, olivine, phlogopite, and serpentinite. minor calc-silicate rocks are layered, with variable grain size, and comprise calcite, phlogopite, quartz, tremolite and locally actinolite. pelitic rocks in this sequence are commonly coarse grained and comprise quartz, biotite, garnet, plagioclase, and sillimanite. large lumps of fibrous sillimanite (up to 3 cm in diameter) are likely pseudomorphs after andalusite. rare pseudomorphs of sillimanite after kyanite were also found. 15 this may indicate variability in pressure conditions in the kangilinaaq belt or prograde barrovian-style metamorphism. quartzo-feldspathic metasedimentary rocks in the central/northern part of the synform are coarse grained, heterogeneous, and rarely garnet-bearing. the qasigiannguit sequence comprises mafic and felsic metavolcanic rocks intercalated with clastic sequences and isolated layers and lenses of strongly deformed, zebrastriped anorthosite. the sequence is c. 700 m wide, trends sw–ne, and is well exposed on the islands south-west of qasigiannguit. the rocks are isoclinally folded with an intrafolially folded gneissic fabric. as mentioned above it is separated from the predominantly clastic sequence of kangilinaaq by 200–500 m of high-strain archaean orthogneiss (fig. 2b; see also below). in contrast with the main supracrustal belt, the metamorphic grade is lower amphibolite facies. the fineto medium-grained, layered amphibolite contains hornblende and plagioclase, and minor clinopyroxene, epidote, biotite, quartz, and possibly also chlorite. only along laksebugt is the amphibolite locally garnet-bearing. felsic layers can be up to several metres wide and contain predominantly plagioclase and quartz, with minor amphibole, white mica and titanite. pelitic and semipelitic layers up to 50 m wide occur mainly on the islands south-west of qasigiannguit. these layers are generally schistose and contain predominantly quartz, plagioclase, biotite and minor garnet, while sillimanite and muscovite are rare. in exposures 25 km north-east of qasigiannguit, an outcrop of kyanite-bearing pelite shows no indications of replacement by sillimanite. this is the only known occurrence in the north-eastern part of the nagssugtoqidian orogen of stable kyanite, although this mineral has also been described from archaean supracrustal rocks within the southern part of the rinkian fold belt (garde & steenfelt 1999). fine-grained, dark grey biotiterich schist/gneiss forms layers 50–80 m wide that grade locally into layered amphibolites. homogeneous, mediumgrained, greyish green, quartzo-feldspathic gneisses form layers up to 30 m wide that are generally platy and contain quartz, plagioclase, white mica, and amphibole. their origin is uncertain. their occurrence in a layered supracrustal sequence, without obvious intrusive contacts, could indicate that these rocks are also of supracrustal origin, but similar rocks in the main kangilinaaq sequence grade into low-strain megacrystic granodiorite to quartz diorite. the contacts between the qasigiannguit amphibolite and the regional archaean orthogneiss are always tectonised, and their original contact relationships are uncertain. however, lenses of amphibolite, and locally also anorthosite, occur abundantly as inclusions in the regional archaean orthogneiss close to its contacts with the supracrustal sequence. these lenses are unlikely to be tectonic because they occur in an irregular pattern, not along zones of high strain. it is more likely that they are xenoliths, suggesting that the orthogneiss precursors intruded into the supracrustal sequence. aeromagnetic characteristics aeromagnetic data covering the ikamiut and kangilinaaq regions (thorning 1993) allow us to image geological features in terms of magnetic responses, also in areas covered by lakes, sea and overburden. a spacing of flight lines of 1 km and a survey altitude of 500 m control the resolution of the aeromagnetic data. in order to enhance the anomaly patterns from shallow-seated geological features, a separation filter has been applied (jacobsen 1987). the filter enhances magnetic anomalies caused by geological features within a specific depth interval in the crust. the rationale of the filter is that the upward continuation of a potential field to a selected height represents the field from sources in the crust below half the selected height. the difference, or residual, between fields at two different heights can then be viewed as representing the field from sources within the corresponding depth interval in the crust. thus, a total magnetic field that has been continued upward to a height of 2 km represents sources in the crust below 1 km. likewise, the field observed at 500 m represents sources below 250 m. consequently, the residual obtained from these two fields by subtraction represents an enhanced image of the anomaly pattern of geological features in the depth interval 250 m – 1 km. the resulting ‘subsurface’ total magnetic field for the ikamiut– kangilinaaq region is shown in figs 3–5. in view of the fundamental ambiguity and complexity of the magnetic field separation, the filtering should only be used as a tool for detection of anomalies and discrimination of patterns, and qualitative interpretations should be supported by other types of geological data. aeromagnetic patterns in the ikamiut region the supracrustal rocks in the ikamiut region appear as intermediate to low magnetic anomalies in the separationfiltered total magnetic field intensity map (–1 to –30 nt, a’s in fig. 4). the orthogneisses are expressed as slightly higher magnetic anomalies (5–20 nt, b and c in fig. 4), although these show considerable variability that may be a consequence of differing contributions from other subsurface lithologies. from the aeromagnetic anomalies it is 16 52° 10 km sydostbugten qasigiannguit kang ilin aaq figs 2b, 5 figs 2a, 4 ikamiut nivaap paa 51° 68°45' 68°30' p p p q o o o h h h o 779 193 139 107 78 55 35 19 4 –10 –22 –32 –42 –51 –59 –67 –74 –81 –87 –93 –98 –104 –109 –114 –118 –122 –127 –131 –136 –140 –145 –151 –157 –163 –170 –179 –189 –204 –226 –359 nt fig. 3. total magnetic field intensity for the ikamiut and kangilinaaq regions. the labels h, o, p and q are referred to in the main text. a shadowing effect from nw (315°n) with an inclination of 45° has been applied. the areas of figs 4 and 5 are outlined by white frames. 68°42' 68°39' 68°36' 52°15' 51°45'52° 2.5 km nt ikamiut nivaap paa a a b g a a d c e f i –36 –27 –20 –16 –13 –10 14 17 20 25 33 60 6 4 2 –1 –3 –7 11 8 fig. 4. the separation-filtered total magnetic field intensity in the interval from 0.250 m – 1 km for the ikamiut area. shadowing effect as in fig. 3. the labels a–i are referred to in the main text. 17 possible to recognise folding of the orthogneiss (e.g. east of c, fig. 4), which correlates closely with the geological mapping (fig. 2a). similarly, a strong magnetic low visible between b and d (fig. 4) defines a refolded fold structure in pelitic gneiss identified during mapping. exposed granitic rocks are visible as high positive anomalies (d, fig. 4). similar anomalies are visible beneath the sea north and east of ikamiut and beneath the bay of nivaap paa (e, f and g, fig. 4). the orthogneiss north of nivaap paa (i, fig. 4) shows slightly lower magnetic anomalies than the orthogneisses to the south. aeromagnetic patterns in the kangilinaaq region the supracrustal belt on the kangilinaaq peninsula appears as a distinct, rather homogeneous, low negative magnetic anomaly (–35 to –300 nt, fig. 5), possibly reflecting the dominance of quartzo-feldspathic rocks with low magnetite contents. small positive, short-wavelength, ovoid anomalies within the supracrustal rocks reflect small ultramafic bodies (j’s, fig. 5, too small to appear on the map of fig. 2b). the biotite-bearing orthogneiss south of the supracrustal belt shows a positive magnetic anomaly (0–55 nt, k in fig. 5). the change is rather abrupt and correlates with the ene-trending mylonitic shear zone along the south-eastern margin of the supracrustal rocks close to the coast of the peninsula (see also fig. 2b). the north-western boundary of the supracrustal belt on the opposite side of the peninsula, where supracrustal rocks and orthogneiss are interleaved, is less well defined. farther north in the vicinity of qasigiannguit, an abrupt change to a high positive magnetic anomaly field to the north-west corresponds with the northern contact of the amphibolite sequence against the orthogneiss (l in fig. n l j j k m 68°55' 51°15' 51° 50°45' 68°50' 68°45' 5 km nt qasigiannguit kang ers une q kang ilin aaq –307 –79 –67 –56 –49 –44 –40 –37 –33 –30 –27 –23 –20 –17 –15 –12 –10 –7 –4 –1 2 6 10 14 17 21 24 26 29 32 36 40 44 48 53 58 65 75 94 162 fig. 5. the separation-filtered total magnetic field intensity in the interval from 0.250 m – 1 km for the kangilinaaq area. shadowing effect as in fig. 3. the labels j–n are referred to in the main text. 18 5). the positive magnetic anomalies within the orthogneiss domains immediately north and south of the kangilinaaq peninsula (m and n in fig. 5) are much stronger than found within the orthogneiss in the north-eastern region of fig. 5 (e.g. at 68°55′n, 50°45′w). based on similar strong positive anomalies associated with granitic rocks in the ikamiut area (fig. 4) and the porphyritic granite on the south-western tip of the kangilinaaq peninsula (fig. 5), the former anomalies may correspond to large deepseated granitic intrusives. likewise, the large magnetic high at eastern sydostbugten may represent granitic intrusive rocks hidden 1–2 km below the present surface (h’s, fig. 3; thorning 1993; nielsen & rasmussen 2004). geochronology zircon separates from three samples from the ikamiut region were analysed to determine (1) the age of emplacement of the regional tonalitic to granodioritic orthogneiss, (2) the age distribution, provenance and minimum age of deposition of the sedimentary precursor to the ikamiut metasedimentary rocks, and (3) the timing of metamorphism. samples of two granodioritic orthogneisses and a quartzo-feldspathic metasedimentary rock were analysed, see below. sample descriptions of these and other rocks are given in table 1, and the age data are presented in tables 2–3 and figs 6–7. all age data in the text are quoted with 2σ absolute uncertainty. methodology samples were crushed and sieved to < 400 µm. the fraction < 45 µm was removed via washing and sieving and the remaining sample panned in water to concentrate the heavy fraction. the heavy, nonmagnetic fraction was separated using heavy liquids (3.30 gcm–3) and a frantz magnetic separator. zircons were hand-picked and mounted in epoxy resin. for secondary ion mass spectrometry (sims) analysis, grains were mounted together with 1065 ma zircons from reference sample 91500, ontario, canada (wiedenbeck et al. 1995). for laser inductively-coupled plasma mass spectrometry (la-icp-ms) analysis no zircon standard was used. the mounted samples were ground to expose the mid-sections of the zircons and polished. the polished samples were examined using backscattered electron (bse) imaging on a philips xl 40 scanning electron microscope at geus, operating at 20kv and a working distance of 10 mm. backscattered electron sample region rock type paragenesis amphibolite 467403 kangilinaaq fine-grained, layered amphibolite 467405 kangilinaaq fine-grained, layered amphibolite 467413 kangilinaaq fine-grained, layered amphibolite 467426 kangilinaaq fine-grained, layered amphibolite 467436 kangilinaaq fine-grained, layered amphibolite hbl-pl ± cpx ± grt, accessory fe-ti oxides 467440 kangilinaaq fine-grained, layered amphibolite 467444 kangilinaaq fine-grained, layered amphibolite 467445 kangilinaaq fine-grained, layered amphibolite 467446 kangilinaaq fine-grained, layered amphibolite orthogneiss 440938 ikamiut medium-grained granodioritic orthogneiss pl-qtz-ksp-bi 467526 ikamiut medium-grained granodioritic orthogneiss pl-qtz-ksp-bi 467401 kangilinaaq medium-grained tonalitic orthogneiss pl-qtz-bi metasedimentary rocks 440910 ikamiut quartzo-feldspathic gneiss pl-qtz-grt-bi, minor sill-ksp 440931 ikamiut quartzo-feldspathic gneiss pl-qtz-bi, minor ksp-mag. bi partly replaced by chl 467503 ikamiut quartzo-feldspathic gneiss pl-qtz-bi, minor ksp-mag-mu 467404 kangilinaaq garnet-bearing schist qtz-pl-bi-sill-grt. bi partly replaced by sill 467408 kangilinaaq cummingtonite gneiss qtz-pl-bi-cu, minor hbl 467417 kangilinaaq biotite-hornblende gneiss qtz-pl-bi-hbl 467420 kangilinaaq fine-grained schist qtz-pl-bi-ep, accessory al & zn 467423 kangilinaaq hornblende gneiss hbl-pl, minor qtz, bi 467433 kangilinaaq hornblende gneiss hbl-pl, minor qtz, bi mineral abbreviations: al: allanite, bi: biotite, cpx: clinopyroxene, cu: cummingtonite, ep: epidote, grt: garnet, hbl: hornblende, ksp: k-feldspar, mag: magnetite, mu: muscovite, pl: plagioclase, qtz: quartz, sill: sillimanite, ti: titanite, zn: zircon. table 1. sample descriptions 19 440938 granodioritic orthogneiss 1 44913 0.207628 2847 4.8 2 54464 0.206401 2843 6.1 3 51744 0.211202 2865 10.9 5 66847 0.205833 2818 12.1 6 45663 0.239096 3072 7.3 7 114515 0.203027 2811 4.9 10 29662 0.206933 2835 8.2 12 81235 0.200267 2785 5.6 15 63304 0.204663 2825 8.7 16 58962 0.203500 2813 7.4 17 32961 0.211432 2888 7.4 20 86073 0.205326 2830 9.7 21 125993 0.199600 2796 4.5 25 60136 0.185466 2664 6.1 26 48587 0.207014 2855 5.6 27 98098 0.202412 2838 8.0 28 46100 0.198084 2836 11.7 29 44476 0.209520 2878 6.4 30 55984 0.205825 2836 9.9 31 52515 0.204633 2835 7.3 33 60481 0.205708 2838 7.2 34 41693 0.217275 2967 29.6 40 16473 0.210866 2905 33.5 41 63641 0.205836 2847 6.9 42 26576 0.199429 2802 10.4 43 34970 0.210589 2882 6.9 44 66299 0.200725 2807 4.9 47 76450 0.185001 2778 5.5 48 344404 0.196981 2819 6.9 50 94026 0.206532 2859 4.6 54 32535 0.206805 2838 17.7 55 18887 0.185070 2753 19.3 57 45001 0.201119 2811 5.3 59 63523 0.203583 2830 8.1 60 35435 0.205041 2842 6.2 61 23531 0.190093 2715 6.2 63 112296 0.208790 2878 4.6 64 39902 0.195071 2763 6.6 74 19592 0.202403 2820 6.2 75 30652 0.204612 2837 11.0 77 41961 0.195999 2774 6.8 78 64028 0.206406 2860 6.1 80 30243 0.209704 2877 6.2 81 106313 0.203552 2844 6.4 82 47863 0.206121 2857 7.4 83 58886 0.201354 2817 6.6 84 119357 0.209519 2882 4.0 86 93017 0.206130 2857 4.8 87 59542 0.208843 2878 4.8 88 80384 0.190889 2738 6.0 89 56425 0.198658 2797 6.6 90 52868 0.194667 2773 5.8 36 30037 0.201451 2808 6.6 37 31254 0.193787 2730 7.6 38 22041 0.205579 2827 10.1 39 52491 0.162084 2504 10.1 40 40780 0.194811 2741 5.7 42 61409 0.195555 2747 4.9 43 103057 0.171459 2534 10.0 44 40828 0.200096 2744 7.8 45 58477 0.197715 2764 5.3 46 20270 0.206049 2843 8.7 47 33049 0.196133 2748 6.5 48 28126 0.193760 2738 6.3 50 19721 0.207732 2845 6.6 51 86891 0.232092 2997 11.6 52 58228 0.201806 2799 5.0 53 50522 0.190800 2710 5.8 54 60856 0.195334 2736 5.6 55 35515 0.198275 2772 5.4 56 51577 0.196653 2758 5.9 58 27007 0.203361 2811 5.7 59 49414 0.202192 2798 5.8 62 10375 0.207836 2855 9.5 63 120563 0.189864 2717 4.2 64 23252 0.187771 2722 7.0 65 26567 0.202492 2801 6.5 69 89976 0.191213 2728 4.7 70 25130 0.199809 2796 6.1 71 68881 0.206577 2861 4.6 72 53407 0.199076 2805 5.1 73 51707 0.194570 2755 4.9 74 67393 0.179579 2625 5.3 75 30132 0.198104 2773 6.6 76 36875 0.197163 2802 7.8 77 28362 0.198893 2790 5.4 78 27133 0.201214 2812 5.7 79 133087 0.188329 2702 4.7 80 45162 0.206677 2856 4.4 81 51961 0.194173 2749 4.6 83 27444 0.201457 2807 5.6 84 26057 0.208900 2868 5.4 85 31842 0.198490 2789 7.0 86 114220 0.180764 2635 5.5 87 62748 0.189371 2717 5.5 88 42113 0.198204 2782 5.2 89 24400 0.199600 2793 6.6 90 29322 0.208314 2861 5.3 91 75430 0.205029 2835 6.0 93 24848 0.198066 2773 6.3 95 60703 0.191617 2717 6.7 96 40762 0.207364 2850 4.1 97 131809 0.193200 2737 6.3 98 43657 0.197222 2778 5.6 99 17404 0.199360 2767 19.2 91 54794 0.203058 2825 6.2 92 86828 0.211233 2895 4.7 93 86433 0.207258 2863 4.4 94 41347 0.203007 2830 5.5 467526 granodioritic orthogneiss 5 56698 0.200720 2804 5.3 6 33787 0.193076 2748 6.6 7 55326 0.195289 2762 4.8 8 76699 0.182851 2652 4.9 10 31196 0.183045 2654 6.6 11 4904 0.204104 2807 12.9 13 39648 0.191787 2718 5.5 15 40707 0.202256 2816 5.0 17 29635 0.217163 2844 19.1 18 6853 0.206028 2827 14.9 19 9339 0.206467 2995 5.7 440910 metasedimentary rock 1 35535 0.200401 2828 6.5 2 31197 0.211066 2824 10.1 3 41461 0.193703 2741 6.1 4 30241 0.204419 2821 6.0 5 25534 0.191383 2718 5.6 6 46366 0.198531 2774 6.1 7 86549 0.207000 2846 5.7 8 48635 0.201767 2807 5.3 9 36751 0.209132 2858 7.0 10 27744 0.183921 2654 9.9 11 43029 0.194263 2742 6.7 12 29314 0.172110 2520 8.6 13 47160 0.206249 2841 5.9 14 49306 0.201659 2804 7.0 15 29978 0.198089 2779 6.0 17 53134 0.194413 2748 5.4 18 27856 0.201215 2810 6.9 19 30082 0.205162 2842 5.9 20 55684 0.198413 2780 5.1 21 147594 0.191303 2722 5.3 22 25436 0.210142 2870 7.0 23 29257 0.196016 2766 7.0 24 20655 0.203877 2826 6.4 25 26159 0.201727 2796 6.6 26 82319 0.215426 2917 4.5 27 112103 0.194672 2751 5.1 30 15491 0.200465 2794 7.1 31 21709 0.204749 2831 8.0 32 14364 0.206770 2843 11.0 33 41967 0.201274 2800 6.8 34 56320 0.189864 2710 5.6 35 23757 0.202176 2801 9.1 table 2. zircon la-icp-ms 207pb-206pb data spot 206pb (cps) 207pb/206pb age (ma) 2σ %spot 206pb (cps) 207pb/206pb age (ma) 2σ % spot 206pb (cps) 207pb/206pb age (ma) 2σ % 20 (bse) images of some of the analysed grains showing sites of analysis and ages obtained are presented in fig. 6. all three samples were analysed at geus using a perkinelmer 6100 drc quadrupole inductively-coupled plasma mass spectrometer combined with a cetac lsx 200 laser ablation unit based on a solid-state nd-yag laser, emitting at a wavelength of 266 nm. the laser was operated at 20 hz with a spot size of 30 µm, producing pits of c. 50 µm depth. the masses 208pb, 207pb, 206pb, and 204pb were analysed in line scans run at 1 µm per second. each analysis comprised 150 time-resolved replicates (duration of total analysis 150 s). in the case of small grains with diameters < 100 µm, only 100 replicates over 100 s were collected. if inconsistencies in the measured ratios were identified within the time span of each analysis, such as spikes relating to inclusions, or significant changes in pbpb ratios indicative of sampling of different age zones, then the whole analysis was discarded. the analyses were standardised against nist 610 glass (pearce et al. 1997) to account for instrument drift. the influence of common pb cannot be assessed using this method, since 204pb was generally below the detection limit. also, because u isotopes could not be measured, the significance of pb loss cannot be assessed, and therefore the ages determined should be regarded as minimum ages. in addition, the age resolution on any individual analysis was restricted owing to relatively low count rates obtained. however, it is a great advantage of the la-icp-ms method that a large number of analyses can be made within a short time period, allowing analysis of large numbers of grains in samples with isotopically simple zircon. the la-icp-ms age data are presented in table 2 and as histograms coupled with relative probability curves (fig. 7). sims analysis of zircon from two samples (440938, 440910) was carried out using a cameca ims 1270 secondary ion mass spectrometer at the nordsim laboratory, swedish museum of natural history, stockholm. the polished zircon mounts were coated with a c. 30 nm layer of gold. analytical procedures and common lead corrections are similar to those described by whitehouse et al. (1997). a primary o2– ion beam is focussed into a spot with a diameter of 20 µm that sputters material from the sample to leave a flat-bottomed crater. positive ions sputtered from the crater are extracted and mass-separated into the peaks of interest: 90zr2 16o, 204pb, 206pb, 207pb, 208pb, 238u, 232th16o, and 238u16o. calibrations of pb/u ratios are based on the observed relationship between pb/u and uo2/u. weighted average 207pb/206pb ages were calculated using isoplot (ludwig 2000). sims age data are presented in table 3 and on tera-wasserburg diagrams in fig. 7. 440938 granodioritic orthogneiss 3 155 94 113 0.607 0.04 0.19819 0.28 14.5789 1.06 0.53352 1.03 –2.4 2811.3 4.6 4 575 274 417 0.502 0.07 0.19781 0.18 14.8262 1.04 0.54361 1.03 –0.4 2808.2 2.9 8 192 24 127 0.117 0.05 0.19575 0.28 14.5329 1.07 0.53845 1.03 –0.6 2791.1 4.6 14 164 23 109 0.140 0.02 0.19767 0.38 14.7352 1.09 0.54065 1.03 –0.9 2807.1 6.1 20a 206 56 134 0.222 0.09 0.19216 0.30 13.6935 1.07 0.51684 1.03 –3.3 2760.7 5.0 24 175 37 117 0.170 0.05 0.19891 0.31 14.7222 1.07 0.53682 1.03 –2.1 2817.2 5.0 29 578 153 384 0.217 0.23 0.19498 0.17 14.2171 1.04 0.52885 1.03 –2.1 2784.6 2.8 41 632 35 391 0.054 0.01 0.19138 0.14 13.5355 1.04 0.51295 1.03 –3.8 2754.1 2.4 54 273 155 199 0.555 0.09 0.19767 0.29 14.6540 1.07 0.53768 1.03 –1.5 2807.0 4.8 66 251 227 191 0.878 0.07 0.19836 0.23 14.3381 1.05 0.52425 1.03 –4.2 2812.7 3.7 72 203 158 150 0.762 0.09 0.19789 0.29 14.2560 1.07 0.52248 1.03 –4.3 2808.9 4.7 76 157 49 107 0.307 0.02 0.20014 0.33 14.6609 1.09 0.53127 1.04 –3.5 2827.4 5.4 440910 metasedimentary rock 22 850 51 526 0.057 0.01 0.18523 0.20 13.1609 1.78 0.51532 1.77 –0.9 2700.3 3.4 29 480 2 308 0.002 0.03 0.18986 0.27 14.0987 1.79 0.53857 1.77 1.6 2740.9 4.4 42 609 3 390 0.003 0.05 0.18800 0.22 13.9521 1.78 0.53823 1.77 2.3 2724.8 3.6 54 554 2 344 0.004 0.01 0.18507 0.22 13.3561 1.78 0.52342 1.77 0.7 2698.8 3.6 58c 106 32 71 0.275 0.23 0.19043 0.43 14.0439 1.82 0.53488 1.77 0.7 2745.9 7.1 58r 610 2 378 0.003 0.04 0.18916 0.21 13.6001 1.79 0.52146 1.77 –1.3 2734.8 3.5 82 1178 7 730 0.004 0.07 0.18473 0.17 13.3129 1.78 0.52267 1.77 0.7 2695.8 2.9 table 3. zircon ion probe (sims) u-th-pb data spot u th pb th/u f 206 % 207pb σ % 207pb σ % 206pb σ % discordance % 207pb σ ppm ppm ppm measured 206pb 235u 238u (conventional) 206pb ages (ma) errors on ratios and ages are quoted at the 1σ level. c: core; r: rim; f 206 %: the fraction of common 206pb, estimated from the measured 204pb. discordance %: degree of discordance of the zircon analysis (at the centre of the error ellipse). 21 orthogneiss the two samples of granodioritic orthogneiss selected for geochronology (440938 and 467526) were collected at the south coast of langesund (fig. 2a). sample descriptions and chemical composition are presented in tables 1 and 4. sample 440938 yielded abundant zircon and contains common thin, transposed granitic layers, whereas sample 467526 does not contain such granitic leucosome. the zircons from both samples are 100–600 µm (mostly c. 200–300 µm) in length and translucent with a heterogeneous orange colour. the crystals are euhedral with slightly rounded terminations and aspect ratios from 1:2.5–1:4, typically c. 1:3. broad oscillatory zones c. 10– 30 µm wide are very common, with rare development of bright, presumably metamorphic rims (see below). the zircons are commonly weakly to moderately fractured, both concentrically and radially, and often show partial fracture healing within bright oscillatory zones (fig. 6). la-icp-ms analyses of 57 oscillatory zoned grains from sample 440938 give a weighted mean age of 2831 ± 23 ma (2σ, mswd = 0.36; table 2; fig. 7a). ten sims analyses of cores of oscillatory zoned grains and two of bright rims reveal more age complexity, with seven of the cores yielding an important 2820–2810 ma age component (table 3; fig. 7b, black data ellipses). four of these seven analyses lie slightly off concordia. another, slightly older and likewise discordant grain (2827 ma, blue in fig. 7b) belongs to the 2831 ± 23 ma la-icp-ms age group. many of the analysed grains show slight discordance indicating partial pb loss, the timing of which is unclear from the data available. two bright rims with significantly younger ages of 2761 ± 10 ma and 2754 ± 5 ma are interpreted as metamorphic. this is supported by the very low th/u of the latter (0.054), though the former is not anomalous in this respect (th/u = 0.22). the sims data are slightly but significantly younger than the la-icp-ms data for the same sample (but from different analysed grains), although the 2σ error on the la-icp-ms age spectrum encompasses most of the sims data. it is difficult to establish the reason for this, particularly given the apparent complexity of the zircons (fig. 7b). it is possible that a larger proportion of older material has been sampled in the la-icp-ms work. it may also be that matrix effects had some influence in standardising zircon data against nist610 glass, though such effects are not generally regarded as significant. the few zircons separated from sample 467526 were analysed via la-icp-ms. eleven analyses of cores displaying oscillatory zonation yield a poorly constrained weighted mean age of 2741 ± 53 ma (2σ, mswd = 0.55; fig. 7c). it might be considered that the large analytical error for this sample leaves room for age complexity, possibly involving analysis of both inherited grains and pb loss (similar to sample 440938). however, all analyses statistically belong to the same population, and there are no significant differences in internal zircon morphology that might account for different age groups. metasedimentary rocks sample 440910 from the ikamiut belt (fig. 2a) is a medium-grained, garnet-bearing quartzo-feldspathic, gneissic rock. colourless, pale yellow and pale pink zircon grains are abundant. they are elongate and generally 100–200 2817 ± 10 200 µm 200 µm b 440938 440910 2700 ± 7 2699 ± 7 2696 ± 6 2746 ± 14 2811 ± 10 2807 ± 10 a fig. 6. backscattered electron images of zircons in samples 440938 (orthogneiss) and 440910 (metasediment) analysed via sims, showing analysed areas and ages obtained. 22 µm in length, with rounded terminations and aspect ratios of 1:1–1:3, typically c. 1:2. bse imaging reveals relatively wide (c. 10–30 µm) oscillatory zoned cores with moderate to well-developed bright rims c. 10–60 µm wide, which in many cases have annealed former fractures (fig. 6). 207pb-206pb ages from 87 la-icp-ms analyses of oscillatory zoned zircon cores are shown in fig. 7d. the vast majority (79) define a tightly clustered peak at c. 2800 ma. the complete age range spans 2997–2520 ma. the youngest ages (< 2700 ma) may reflect mixed core–rim data, although most of them are statistically within the main age population. if a few anomalous old and young ages are disregarded, a weighted mean age of 2779 ± 18 ma is obtained (2σ, mswd = 0.32, n = 79). this group is interpreted as comprising a homogeneous population of detrital zircons, consistent with local derivation from orthogneiss of this age (see above). seven sims analyses, of one core with oscillatory zonation and six bright rims, all fall on concordia (fig. 7e). the core gives the oldest age of 2746 ± 14 ma (th/u = 0.275). the six analyses of bright rims give ages between 2741 ± 9 and 2696 ± 6 ma. all have very low th/u (0.004–0.06), consistent with a metamorphic origin. these metamorphic ages are comparable to the few young ages also identified in the laicp-ms analyses, and suggest that this sample underwent metamorphism at c. 2740–2700 ma, shortly after its deposition. an alternative, and in our view less likely interpretation is that the metamorphic rims were developed during metamorphism of the source rock prior to erosion and deposition of the sediment. age (ma) 2779 ± 18 ma (2σ) mswd = 0.32 n = 79 440910 467526 2741 ± 53 ma (2σ) mswd = 0.55 n = 11 age (ma) 440938 orthogneiss d c a 2831 ± 23 ma (2σ) mswd = 0.36 n = 57 age (ma) 16 14 12 10 8 6 4 2 0 16 14 12 10 8 6 4 2 0 6 4 2 3 0 5 1 n um be r of a na ly se s n um be r of a na ly se s n um be r of a na ly se s 1500 2000 2500 3000 3500 1500 2000 2500 3000 3500 1500 2000 2500 3000 3500 orthogneiss metasedimentary rock 20 7 p b/ 20 6 p b 20 7 p b/ 20 6 p b 238u/206pb 238u/206pb 2770 2750 2730 2710 2690 0.182 0.184 0.186 0.188 0.190 0.192 0.194 1.7 1.8 1.9 2.0 2.1 440910e 0.188 0.192 0.196 0.200 0.204 440938b 2840 2820 2800 2780 2760 2740 1.80 1.9 2.0 fig. 7. la-icp-ms 207pb/206pb histograms and sims u-pb tera-wasserburg concordia plots of zircon age data from the ikamiut region. a: histogram of 207pb/206pb ages from orthogneiss sample 440938. b: terawasserburg plot for orthogneiss sample 440938 showing nine cores of oscillatoryzoned grains (black), an older core (blue), and two bright rims (red). c, d: histograms of 207pb/206pb ages from orthogneiss 467526 and metasedimentary sample 440910. e: tera-wasserburg plot for metasedimentary sample 440910 showing one core (black) and six metamorphic rims (red). the histogram bin size is 25 ma (a, d) or 50 ma (c). error ellipses on concordia diagrams are drawn at 68.3% confidence (1σ). 23 geochemistry whole-rock geochemical analysis of amphibolites, felsic orthogneiss and metasedimentary rocks from the ikamiut and kangilinaaq regions was undertaken to (a) geochemically characterise these rock types, (b) investigate the likely tectonic environment of formation and provenance of the amphibolites and metasedimentary rocks, and (c) investigate likely regional correlations. sample descriptions are presented in table 1, and major and trace element compositions in table 4. analytical procedure major and trace element analyses (table 4) were performed by geus. the samples were ground in tungsten (ikamiut samples) or agate (kangilinaaq samples) mills, and dried. for major elements the rock powders were fluxed with sodium tetraborate and fused to glass discs and analysed with a philips pw1606 x-ray fluorescence (xrf) mass spectrometer. na and cu were determined by atomic absorption spectrometry, and volatiles were analysed by gravimetry. refer to kystol & larsen (1999) for the complete analytical procedure. for trace element analyses, powdered samples were brought into solution and analysed using a perkinelmer 6100 drc quadrupole icpms instrument. for the ikamiut samples zr, cr, rees and hf were determined by dissolving a piece of the borate glass used in the major element xrf analyses, in order to obtain complete contributions of these elements from chromite and zircon. the icp-ms results were corrected for the relevant oxide interferences using bhvo-1 and gh as standards. for the kangilinaaq samples some trace elements were also analysed by xrf performed directly on pressed powder tablets at the geological institute, university of copenhagen, using a phillips pw 1400 xrf spectrometer. the data were corrected for matrix variations using the major element compositions, and agv-1 was run as standard. orthogneisses two orthogneiss samples from the ikamiut region were analysed (440938 and 467526), and one from kangilinaaq (467401; table 4a; fig. 2). the two ikamiut samples show very similar granodioritic major element chemistry, while the kangilinaaq sample is more tonalitic. the kangilinaaq sample has low ree concentrations and a fairly steep ree curve with la n /lu n ~ 16 (fig. 8b). in contrast, the two ikamiut samples have higher ree concentrations and significant negative eu anomalies, consistent with the more evolved composition of these rocks. amphibolites from the kangilinaaq belt nine amphibolites s.s. from the kangilinaaq belt were analysed for major and trace elements (table 4a); no amphibolite samples have been analysed from the ikamiut region, where amphibolites only constitute a minor component of the supracrustal rocks. the nine samples from kangilinaaq show only a small range in chemical composition. they have relatively primitive signatures with low sio 2 (46–49 wt%) and high mgo (7–11 wt%), and flat ree patterns that group tightly around ten times chondrite values (fig. 8a). their ti/v ratios display a narrow range of 16–19, and in a ti-v diagram (shervais 1982) they plot just within the island arc field (fig. 9). the positive correlation between ti and v could reflect fractionation of olivine and plagioclase. the geochemical resemblance between all nine samples and their well-defined ti/v trend are consistent with formation within a single volcanic suite. metasedimentary rocks and hornblendebearing gneisses from the kangilinaaq belt metasedimentary sample 467404 is characterised by high alumina (19 wt%) coupled to low cao (0.7 wt%) and high concentrations of ree and ba (605 ppm), consistent with a clay-rich precursor. the ree curve is steep (la n /lu n = 30, la = 140 times chondrite), and has a significantly negative eu anomaly. sample 467420 is siliceous (75 wt% sio 2 ), consistent with a relatively mature sedimentary precursor. its ree concentrations lie just below those of sample 467404, with a similar steep ree pattern. for sample 467417, both major and trace elements agree well with the average composition of archaean mudstone from taylor & mclennan (1985, table 7.8). the ree curve resembles that of sample 467408 (see below), although it has slightly lower concentrations of the hree. the hornblende-bearing gneisses 467423 and 467433 were collected from thin (< 1 m) amphibolite units within metasedimentary sequences. the geochemical compositions of these two samples are close to those of the nine amphibolite samples described above, and they can only be distinguished from the latter by their higher concentrations of lree, ba, and sr, and higher k 2 o and rb in 24 467403 467405 467413 467426 467436 467440 467444 467445 467446 440938 467526 467401 sio2 48.73 48.72 49.3 46.86 48.39 47.37 48.01 48.25 46.00 71.93 71.70 68.86 tio2 1.03 0.87 0.83 0.60 0.64 0.66 0.90 0.88 0.75 0.24 0.39 0.22 al2o3 13.16 15.81 14.61 16.21 15.68 16.36 13.29 12.18 16.25 14.58 14.64 16.93 fe2o3 1.67 2.54 2.7 1.88 2.13 2.79 1.90 1.88 2.83 1.99 2.09 0.09 feo 10.64 8.89 9.11 7.48 7.85 8.14 8.04 8.82 7.94 0.00 0.00 1.68 feo* 12.14 11.17 11.53 9.17 9.77 10.65 9.75 10.51 10.49 1.79 1.88 1.76 mno 0.22 0.21 0.20 0.18 0.17 0.16 0.13 0.19 0.17 0.01 0.02 0.02 mgo 8.82 6.90 7.73 8.30 8.34 9.37 10.74 11.15 10.04 0.57 0.57 0.98 cao 11.46 12.72 11.51 14.17 12.78 10.91 12.31 12.43 10.42 2.02 2.12 4.02 na2o 2.25 1.28 1.58 1.60 1.36 2.00 1.97 1.56 2.38 4.47 4.70 5.17 k2o 0.11 0.23 0.25 0.31 0.25 0.08 0.52 0.10 0.48 3.08 2.31 1.05 p2o5 0.06 0.06 0.05 0.04 0.04 0.04 0.06 0.06 0.05 0.07 0.08 0.07 volatiles 1.58 1.5 1.57 1.48 1.57 1.6 1.43 1.53 2.06 0.27 0.10 0.53 sum 99.74 99.72 99.43 99.11 99.21 99.47 99.29 99.02 99.36 99.23 98.71 99.62 sc 53.6 49.0 54.8 44.1 44.2 38.7 49.6 52.3 49.7 5.5 8.1 4.6 v 347 283 308 214 237 220 301 297 256 16 13.8 24.1 cr 106 301 380 444 450 207 521 485 411 5.0 2.6 35.9 co 53.5 55.1 51.6 51.2 56.1 58.0 70.3 56.6 54.0 17.2 14.7 6.0 ni 117 150 160 221 222 254 214 221 159 3.6 3.4 12.9 cu 64.2 89.0 82.7 39.1 90.3 89.5 4.8 82.8 80.8 2.4 8.1 7.3 zn 91.9 86.0 92.2 74.3 74.8 80.1 40.0 77.3 74.6 40.9 39.8 45.2 ga 15.7 16.3 15.7 14.2 14.2 14.8 14.8 14.1 15.4 18.2 20.8 23.2 rb 1.1 7.6 9.1 12.4 10.9 1.2 9.5 2 10.6 82.1 91.4 92.9 sr 124 111 91 118 150 100 94 109 155 385 303 375 y 19.6 19.5 19.6 13.6 14.3 14.2 18.1 17.4 17 3 7.8 2.8 zr 29.7 11.8 12.5 9.4 9.4 15.8 18.4 11.6 16.2 143 117 57.9 nb 3.1 2 1.9 1.5 1.4 1.1 2.4 2.2 1.7 2.5 4.9 15.4 cs 0.0 0.3 0.2 0.7 0.8 0.0 0.1 0.0 0.2 1.2 2.6 5.6 ba 18 56 43 47 25 6 31 20 60 748 446 210 la 3.2 2.2 2.4 1.7 1.6 1.6 3.5 2.5 2.2 34.1 18.9 4.0 ce 8.4 6.0 6.3 4.5 4.3 4.3 8.0 6.5 6.0 65.0 37.8 8.3 pr 1.3 1.0 1.0 0.7 0.7 0.7 1.2 1.0 0.9 7.2 4.2 1.0 nd 6.7 5.3 5.3 3.7 3.7 3.7 6.0 5.6 5.1 24.2 14.9 3.7 sm 2.2 1.8 1.8 1.4 1.4 1.3 1.9 1.9 1.7 3.3 2.5 0.8 eu 0.7 0.7 0.7 0.5 0.5 0.6 0.7 0.7 0.6 0.6 0.6 0.3 gd 3.0 2.7 2.6 1.6 1.7 1.6 2.4 2.3 2.2 3.4 2.6 1.0 tb 0.5 0.5 0.5 0.3 0.3 0.3 0.5 0.4 0.4 0.2 0.3 0.1 dy 3.3 3.2 3.1 2.3 2.3 2.3 3.1 2.9 2.8 0.9 1.5 0.6 ho 0.7 0.7 0.7 0.5 0.5 0.5 0.6 0.6 0.6 0.1 0.3 0.1 er 1.9 1.9 2.0 1.3 1.4 1.4 1.8 1.7 1.6 0.3 0.7 0.2 tm 0.3 0.3 0.3 0.2 0.2 0.2 0.3 0.2 0.2 0.0 0.1 0.0 yb 1.8 1.9 2.0 1.3 1.5 1.4 1.7 1.6 1.6 0.2 0.7 0.2 lu 0.3 0.3 0.3 0.2 0.2 0.2 0.3 0.3 0.3 0.0 0.1 0.0 hf 1.0 0.6 0.6 0.4 0.5 0.6 0.7 0.6 0.6 3.8 3.3 1.5 ta 0.2 0.1 0.1 0.1 0.1 0.1 0.1 0.1 0.1 0.5 1.1 1.9 pb 1.4 1.5 1.1 1.3 0.8 0.5 0.6 1.3 1.2 9.5 7.6 7.4 th 0.4 0.2 0.2 0.2 0.2 0.1 0.3 0.3 0.2 10.1 4.2 0.9 u 0.1 0.1 0.1 0.1 0.1 0.0 0.2 0.1 0.0 0.5 0.9 0.4 total ree 34 28 29 20 20 20 32 28 26 140 85 20 amphibolite, kangilinaaq belt table 4a. chemical analyses of amphibolite and orthogneiss orthogneiss major elements (in wt%) by xrf at geus. trace elements (in ppm) by icp-ms at geus. feo* = total fe calculated as feo. volatiles = loss on ignition corrected for oxygen uptake due to oxidation of iron. 25 440910 440931 467503 467404 467408 467417 467420 467423 467433 sio2 63.27 69.36 74.26 62.76 61.37 59.87 75.29 49.25 49.94 tio2 0.59 0.49 0.03 0.61 0.67 0.64 0.34 0.61 0.78 al2o3 17.61 16.08 14.81 19.09 13.90 16.78 13.01 16.21 15.53 fe2o3 5.19 2.38 0.01 1.17 1.32 1.25 0.51 3.63 1.80 feo 0.00 0.00 0.37 5.82 6.73 5.38 1.26 6.97 8.34 feo* 4.67 2.14 0.38 6.87 7.92 6.50 1.72 10.24 9.96 mno 0.05 0.01 0.00 0.05 0.11 0.15 0.01 0.19 0.16 mgo 2.31 1.03 0.10 2.80 6.41 3.97 0.57 7.70 7.78 cao 2.26 3.17 2.02 0.75 4.60 3.68 5.62 9.21 10.73 na2o 3.84 3.85 4.89 1.60 1.82 2.73 1.05 2.63 1.88 k2o 3.02 1.79 2.81 3.20 1.46 3.21 0.62 1.63 0.50 p2o5 0.13 0.11 0.02 0.03 0.07 0.08 0.05 0.13 0.16 volatiles 1.00 0.75 0.29 1.64 1.26 1.28 0.90 1.39 1.48 sum 99.27 99.02 99.62 99.52 99.71 99.01 99.23 99.55 99.06 sc 18.0 10.2 0.8 17.6 39.3 26.6 7.7 43.8 39.7 v 98.5 38.2 1.7 56.7 208 149 43.5 204 209 cr 89.1 25.0 0.4 120.0 667 312 66.1 268 318 co 22.3 17.2 26.9 16.3 44.8 28.4 14.0 51.2 50.7 ni 32.3 5.2 1.3 54.7 194.0 106.0 34.0 148.0 145.0 cu 21.9 8.8 5.2 4.8 20.0 26.3 26.9 25.1 34.8 zn 71.4 38.9 4.4 37.8 89.5 104.5 41.4 81.8 87.4 ga 20.8 19.6 16.0 25.7 17.5 23.1 15.5 16.3 15.9 rb 82.7 48.2 58.5 96.9 51.1 197.0 41.6 59.1 8.0 sr 210 314 402 47 117 177 137 270 288 y 16.0 6.4 2.4 11.1 15.4 16.7 11.1 17.4 17.0 zr 134 151 70.8 57.7 82.4 83.1 73.9 38.4 24.0 nb 5.6 5.6 0.8 8.7 3.3 6.9 2.9 2.3 2.0 cs 4.1 4.7 0.5 2.0 2.5 8.1 2.7 2.4 0.0 ba 417 286 591 581 386 377 310 153 97 la 28.2 20.0 7.9 34.5 11.0 9.6 28.3 11.9 13.4 ce 58.3 40.6 16.1 64.7 29.1 28.5 56.0 29.5 37.5 pr 6.9 4.4 2.1 7.4 2.9 2.6 6.5 3.8 4.5 nd 25.3 15.7 7.7 26.4 11.8 10.5 24.5 16.8 19.7 sm 4.3 2.6 1.8 4.4 2.7 2.5 4.0 3.9 4.4 eu 1.2 0.7 0.4 0.7 0.7 0.7 1.0 1.2 1.1 gd 4.5 2.6 1.5 5.0 3.6 2.3 3.9 3.5 3.8 tb 0.6 0.3 0.2 0.5 0.5 0.4 0.4 0.5 0.5 dy 3.0 1.4 0.6 2.4 2.8 2.4 2.3 2.9 3.0 ho 0.6 0.2 0.1 0.4 0.6 0.6 0.4 0.6 0.6 er 1.6 0.6 0.2 1.1 1.6 1.9 1.1 1.7 1.7 tm 0.2 0.1 0.0 0.1 0.2 0.3 0.1 0.3 0.3 yb 1.5 0.5 0.2 0.9 1.5 2.1 0.8 1.6 1.6 lu 0.2 0.1 0.0 0.1 0.2 0.3 0.1 0.2 0.2 hf 3.7 3.9 2.8 1.5 2.2 2.2 1.8 1.0 0.8 ta 0.9 2.6 1.9 0.6 0.3 0.6 0.2 0.1 0.1 pb 12.7 5.2 11.8 3.7 4.7 9.5 6.1 6.6 4.2 th 6.2 3.3 4.8 8.5 3.5 6.0 4.2 2.0 2.0 u 1.5 0.9 1.3 0.8 0.9 3.1 0.9 0.7 0.6 total ree 137 90 39 149 69 65 129 79 92 major elements (in wt%) by xrf at geus. trace elements (in ppm) by icp-ms at geus. volatiles: loss on ignition corrected for oxygen uptake due to oxidation of iron. feo*: total fe calculated as feo. table 4b. chemical analyses of various supracrustal rocks 26 sample 467423. sample 467408 has an intermediate silica content (61.37 wt% sio 2 ), is cummingtonite-bearing, and has high concentrations of mgo (6.41 wt%) and feo* (7.92 wt%) as well as cr and n (627 and 186 ppm, respectively). it also has high ba (405 ppm). the ree curve is almost flat, with ten times chondritic hree and a weak lree enrichment (la n /lu n = 5). there is a small positive ce anomaly in addition to a negative eu anomaly. metasedimentary rocks from the ikamiut belt sample 440910 is the most aluminous (18 wt% al 2 o 3 , table 4), has high ree concentrations, and a fairly steep ree curve with la n /lu n = 15 (fig. 8c). samples 440931 and 467503 are more siliceous (69 and 74 wt% sio 2 , respectively) with lower al, fe and mg. both have lower ree concentrations and slightly steeper ree curves than sample 440910. interpretation the metasedimentary rocks and amphibole-bearing gneisses of supracrustal origin described above from the ikamiut and kangilinaaq belts can be divided into two groups based on their geochemical compositions and ree patterns. five of them, namely all three ikamiut samples and samples 467404 and 467420 from kangilinaaq, are typical metasedimentary lithologies with steep ree curves (fig. 8c). although they have varying ree concentrations, all five samples have fairly steep ree curves that are comparable to the ree patterns seen in the kangilinaaq and ikamiut orthogneisses. the group shows a trend of increasing ree concentrations with decreasing sio 2 and increasing al 2 o 3 , consistent with the general presumption that the ree are preferentially concentrated in the clay fraction of sediments. one exception from this is the siliceous sample 467420 that has ree concentrations comparable to the most aluminous metasediments. its unusual ree enrichment may be due to high contents of detrital allanite and zircon, as these minerals incorporate high ree concentrations. the ree curves for the four amphibole-bearing gneisses (467408, 467417, 467423, and 467433; fig. 8d) have la n /lu n ~ 5, showing significantly flatter patterns. these four samples all have high concentrations of mafic minerals and may represent intermediate tuffaceous rocks or mildly chemically altered mafic volcanic rocks. the rather peculiar composition of sample 467408, with high ni and cr, suggests that it is a metamorphosed, hydrothermally 100 10 1 ree sample / chondrite la ce pr nd pm sm eu gd tb dy ho er tm yb lu la ce pr nd pm sm eu gd tb dy ho er tm yb lu 100 10 1 100 10 1 la ce pr nd pm sm eu gd tb dy ho er tm yb lu 100 10 1 la ce pr nd pm sm eu gd tb dy clastic metasedimentary rocks amphibolites amphibole-bearing paragneisses felsic orthogneisses 467503 440931 440910 467404 467420 440938 467526 467401 c b d a 467408 467433 467423 467417 fig. 8. chondrite-normalised ree plots. a: amphibolites. b: felsic orthogneiss. c: quartzo-feldspathic metasedimentary rocks. d: mafic metasedimentary rocks. blue: samples from the kangilinaaq region. red: samples from the ikamiut region. 27 altered mafic volcanic rock. alternatively, such high ni and cr in a clastic sedimentary precursor would require an abundance of heavy minerals such as garnet and spinel. discussion regional structures both the ikamiut and kangilinaaq regions preserve complex tectono-metamorphic histories, and although they show similarities in lithologies and metamorphic grade (amphibolite facies mineral assemblages defining d1 and d2 structures) and lie roughly along strike, the large structures are sufficiently different to make a direct correlation between the two regions and their supracrustal belts difficult and dubious. both regions show evidence for at least two generations of fold structures, with kilometreto ten kilometre-scale f2 folds dominating the outcrop pattern. however, there is considerable variation in the typical strike of foliation and plunge of large-scale folds and lineations between the two regions. furthermore, the aeromagnetic data do not suggest a strong link between the two regions. on the basis of geological mapping and the aeromagnetic data it seems likely that the kangilinaaq belt forms a synformal fold closure at the south-western tip of the kangilinaaq peninsula, with little or no westward continuation. we cannot rule out the possibility that there is continuation of this belt across sydostbugten into the ikamiut region along an abrupt change in the aeromagnetic response across northern sydostbugten (boundary o in fig. 3). this could be interpreted as an extension of the linear aeromagnetic anomaly marking the northern contact of the qasigiannguit amphibolite sequence with the surrounding orthogneiss, which appears to extend westward to several kilometres north of the nivaap paa bay (boundary p in fig. 3). however, there is no strong field evidence to support this, since no comparable amphibolite sequence occurs on land at q (fig. 3). it is possible that this linear aeromagnetic anomaly relates instead to an interpreted granitic body beneath sydostbugten, represented by outcrop on the south-eastern tip of the kangilinaaq peninsula, and on the south-west coast of nivaap paa. magmatism geochronological data show that the magmatic precursors to granodioritic orthogneiss from the ikamiut region were emplaced in the late archaean. sample 440938, with a la-icp-ms pb-pb zircon age of 2831 ± 23 ma u-pb zircon ages of 2820–2810 ma, is significantly older than sample 467526, which was collected from a nearby locality (la-icp-ms pb-pb zircon age = 2741 ± 53 ma). the latter compares well with a homogeneous undeformed granite collected a few kilometres west of the head of nivaap paa, which yielded an upper intercept u-pb zircon age of 2778 +7/–3 ma (connelly & mengel 2000). similarly, a grey tonalitic orthogneiss sampled close to aasiaat yielded a u-pb concordia age of 2727 +36/–22 ma and consistent pb-pb and rb-sr whole-rock ages of 2759 +87/–92 ma and 2752 ± 656 ma respectively (kalsbeek et al. 1987). available data on emplacement ages of the precursors to the tonalitic orthogneiss in the kangilinaaq region suggest these may be slightly older than those in the ikamiut region. kalsbeek & nutman (1996) reported ion probe data for a few zircon grains from a granodioritic to granitic orthogneiss in the kangilinaaq area, which gave an emplacement age between 2900 and 2750 ma. keiding (2004) presented la-icp-ms 207pb/206pb zircon age data for a tonalitic orthogneiss of 2818 ± 1 ma, interpreted as an igneous crystallisation age. 700 600 500 400 300 200 100 0 20 ti/v = 10 20 50 100 ti (ppm) /100 v arc tholeiite morb oib 0 2 4 6 8 10 12 14 16 18 fig. 9. ti/v discrimination diagram (shervais 1982) for nine amphibolite samples from the kangilinaaq peninsula (fig. 2b), illustrating their island arc affinities. note that ti and v are both immobile elements, considered to be stable during hydrothermal alteration and regional metamorphism (e.g. nicollet & andriambololona 1980; mottl 1983). the partition coefficient of v varies with the oxygen fugacity of the magma, whereas the partition coefficient of ti remains unchanged. 28 sedimentation deposition of the sedimentary precursors to the ikamiut and kangilinaaq belts likely occurred in the neoarchaean, as indicated by the metamorphic ages of c. 2800–2700 ma of zircon from both belts. the age of the kangilinaaq belt is further constrained by the 2723 ± 15 ma emplacement age of a two-mica granite, which cross-cuts pelitic metasedimentary rocks of this belt on the south shore of kangersuneq fjord (thrane & connelly 2006, this volume). as regards the ikamiut belt, its detrital zircon ages do not preclude deposition after the archaean. neoarchaean metamorphic rims have been observed on some detrital grains, but these might have formed already during metamorphism of the source and survived during erosion and deposition. however, we consider this possibility unlikely. a neoarchaean depositional age is furthermore in agreement with the rb-sr isotopic data for 12 metasedimentary samples from ikamiut reported by kalsbeek & taylor (1999), which likewise show that their source was archaean. finally, the ikamiut belt has experienced a more complex structural history than the naternaq supracrustal belt of palaeoproterozoic age to its south (østergaard et al. 2002; garde 2004; thrane & connelly 2006, this volume). this likewise points to an archaean age of the ikamiut belt. the depositional sources themselves are constrained by detrital zircon populations. the detrital age spectrum for zircon grains (with igneous zonation) from metasedimentary sample 440910 from the ikamiut region forms a tightly clustered peak at c. 2800 ma, consistent with the age of the older tonalitic orthogneiss in this area. furthermore, the steep ree pattern of this sample mimics that of the granodioritic to tonalitic orthogneiss that dominates the region. thus the sedimentary precursor to this rock was probably derived locally from (and possibly deposited onto) the igneous precursor to the neoarchaean orthogneiss basement. this requires a tectonic environment conducive to rapid erosion of the precursor to the source orthogneiss shortly after its emplacement at c. 2800 ma. similarly, a metasedimentary rock from the kangilinaaq region contains archaean detrital zircon, with a strong peak at c. 2800 ma (thrane & connelly 2006, this volume, 207pb/206pb zircon). however, there is also evidence for a significant older component, not recognised in metasedimentary rocks from the ikamiut region: keiding (2004) reported detrital zircon ages for two metasedimentary samples from the kangilinaaq region with grains as old as 3600 ma, and down to 2500 ma, although the youngest grains (< 2800 ma) were suspected of having suffered lead loss. both samples show a large spread of ages, but neither has a significant neoarchaean component at c. 2800 ma. these data contrast with those of sample 440910 from the ikamiut region, which has a tightly clustered detrital zircon population at c. 2800 ma. this suggests that at least some of the kangilinaaq metasedimentary rocks were derived from different, older, and distal source rocks: the older (> 2900 ma) component may be derived from a presently unexposed region within the nagssugtoqidian orogen or possibly from the lesser-known craton to the north (keiding 2004). a difference in the depositional sources of the ikamiut and kangilinaaq belts is also apparent from the geochemical data. the metasedimentary rocks form two groups based on their ree patterns. the first group (three ikamiut and two kangilinaaq samples) shows steep ree curves (fig. 8c), interpreted as indicative of derivation from a felsic source, based on their similarity with ree patterns seen in the kangilinaaq and ikamiut orthogneisses. the second group (four kangilinaaq samples, fig. 8d) shows flatter patterns, consistent with derivation from a bimodal source, i.e. detritus of both felsic (steep ree patterns) and mafic (flat ree patterns) igneous rocks. the precursors to the amphibole-bearing gneisses are interpreted as volcaniclastic material that may have been mixed with clastic material during deposition or by tectonic interleaving. given the intensity of deformation and paucity of information on the depositional environment(s) we consider it imprudent to establish a single stratigraphicstructural interpretation. in view of the different dominant lithologies of the qasigiannguit amphibolite and the remainder of the kangilinaaq belt, it would be interesting to investigate further whether they represent the same or different settings. the geochemistry of their amphibolite samples fall within the same range, but no metasedimentary rocks associated with the qasigiannguit amphibolite have been analysed, and these may be important for identifying links between the latter unit and the kangilinaaq belt. metamorphism age data for metamorphic zircon from orthogneiss and metasedimentary samples from the ikamiut region indicate an important neoarchaean thermal event. two analyses of metamorphic rims from the 2831 ± 23 ma orthogneiss 440938 yield ages of 2761 ± 10 ma and 2754 ± 5 ma, within error the same as the 2741 ± 53 ma emplacement age of sample 467526 (fig. 7). this may suggest that continued synkinematic emplacement of neoarchaean granitoids at c. 2760–2700 ma resulted in metamorphism 29 of slightly older (c. 2800 ma) crust. this is also supported by u-pb zircon data from the metasedimentary sample 440910. the six sims ages of metamorphic rims fall in two ranges, 2741 ± 9 ma and 2696 ± 6 ma, and similar young ages were identified in the la-icp-ms data (fig. 7). these ages probably relate to the growth of s 1 garnet, biotite, plagioclase, quartz, minor sillimanite and k-feldspar in this and other metasedimentary rocks, indicative of amphibolite facies conditions only shortly after deposition, and predating regional f 2 folding. neoarchaean metamorphism has also been recognised from zircon age data in the kangilinaaq region. keiding (2004) reported c. 2800 and 2760 ma la-icp-ms ages of zircon rims and discrete grains, interpreted as metamorphic in origin, in a 2818 ± 1 ma tonalitic orthogneiss. these ages correlate reasonably well with 2810–2720 ma metamorphic u-pb zircon and monazite ages in 2870–2810 ma orthogneisses from throughout the nagssugtoqidian orogen (connelly & mengel 2000). it is likely that neoarchaean amphibolite facies metamorphism in the ikamiut and kangilinaaq regions was the product of tectonism along a convergent margin (see also connelly & mengel 2000) on the basis of (1) the tonalitic to granodioritic composition of the neoarchaean regional orthogneisses, (2) the apparent island arc geochemical character of amphibolites of the kangilinaaq belt, (3) differences in the ages of sediment sources in the two supracrustal belts, and (4) the rapidity of the cycle of magmatism, erosion, sedimentation, and metamorphism. proterozoic zircon ages are known from the kangilinaaq region. keiding (2004) reported 1920–1820 ma laicp-ms ages of zircon rims and discrete grains in archaean tonalitic orthogneisses, and also reported weighted mean age of 1919 ± 11 ma from three rims of detrital grains in a metasedimentary rock from the kangilinaaq belt. thrane & connelly (2006, this volume) report metamorphic ages of c. 1850 ma for a metasedimentary rock collected on the south shore of kangersuneq fjord, attributed to the peak of regional nagssugtoqidian metamorphism. given the consistency of ene-trending d 2 structures in the kangilinaaq and ikamiut regions with enetrending palaeoproterozoic structures throughout the orogen, these are interpreted as the product of the c. 1850 ma nagssugtoqidian orogenesis. the significance of the slightly older, c. 1920 ma metamorphic age is not clear, but may indicate that part of this region experienced a thermal event prior to the main regional nagssugtoqidian orogenesis. by contrast, no significant indications of palaeoproterozoic resetting are found in our ikamiut data. the slightly discordant zircon data in sample 440938 suggest some pb loss in this sample, although the timing is not clear. similarly, three titanite u-pb analyses of 2778 +7/–3 ma reported by connelly & mengel (2000) from a homogeneous, undeformed granite plot on a discordia line between 2789 ± 100 and 1775 ± 10 ma. this indicates that palaeoproterozoic metamorphic temperatures were too low to completely reset titanite in this region. likewise, no indication of u-pb resetting in zircon was found in the 2727 +36/–22 ma age from a tonalitic gneiss reported by kalsbeek et al. (1987). this contrasts with zircon u-pb analyses of samples from the nordre strømfjord region in the core of the nagssugtoqidian orogen, which experienced significant pb-loss at c. 1850 ma (kalsbeek et al. 1987). conclusions new mapping, geochemical, geochronological and geophysical studies of two supracrustal belts from sydostbugten, southern disko bugt region, west greenland, shed light on the neoarchaean tectonic evolution of the northern nagssugtoqidian basement. the kangilinaaq belt was deposited at c. 2800 ma, whereas the deposition of the ikamiut belt may postdate c. 2740 ma. the geochemical signatures of the majority of metasedimentary samples from the kangilinaaq region show ree patterns indicative of mixed felsic and mafic sources with distal mesoand palaeoarchaean components that are not currently known in situ in this part of west greenland. island-arc geochemical affinities of intercalated amphibolites are consistent with deposition in an arc setting. in contrast, the ikamiut belt was sourced locally from, and deposited onto or proximal to the igneous precursors of neoarchaean granodioritic to tonalitic orthogneisses. this is constrained by (1) the similarity in ree signatures of metasedimentary rocks and local orthogneisses and (2) the zircon emplacement ages of orthogneisses (c. 2820–2810 ma; 2831 ± 23 ma; 2741 ± 53 ma) and detrital zircons in metasediment (2779 ± 18 ma). zircon u-pb data and s 1–2 sillimanite-bearing mineral assemblages (this study and existing data) indicate that c. 2800–2700 ma amphibolite facies metamorphism affected both regions, shortly after the emplacement of the regional orthogneiss precursors and deposition of the supracrustal rocks. the rather rapid cycle of magmatic emplacement, island arc volcanism, erosion and sedimentation, and subsequent amphibolite facies metamorphism is consistent with neoarchaean convergent tectonism at the northern margin of the present nagssugtoqidian orogen. subsequently, both regions underwent palaeoproterozoic regional deformation and lower amphibolite facies 30 metamorphism at c. 1850 ma during the nagssugtoqidian orogenesis, the effects of which control outcrop patterns in both areas. in the ikamiut region, the supracrustal belt defines a broad, shallowly w-plunging antiformal structure with associated kilometre-scale parasitic f 2 folds. s 1 fabrics are folded into metreto kilometre-scale f 2 folds and variably transposed into ene-striking, steeply dipping s 2 fabrics and shallow w-plunging mineral lineations defined by biotite and muscovite. in the kangilinaaq region, the supracrustal belt defines a broad, ne-plunging f 2 fold structure. a pervasive, nestriking, moderately dipping s 2 fabric, defined by mediumto coarse-grained garnet-hornblende-biotite-sillimanitebearing assemblages in pelitic rocks, is folded into f 3 folds, and attests to amphibolite facies metamorphic conditions during deformation. lack of palaeoproterozoic resetting of the zircon u-pb isotopic system in the ikamiut region, cf. the kangilinaaq region, suggests that temperatures were relatively lower in the former region during the nagssugtoqidian orogenesis. acknowledgements mads sylvest christensen, jane gilotti, christian knudsen, stanislaw mazur, mac persson, sandra piazolo and thomas v. rasmussen contributed to field work in 2002– 2003 reported here and part of the geus project archaean and proterozoic crustal evolution in the aasiaat region, central west greenland. dirk frei, mark t. hutchison, lev ilyinsky, jørgen kystol, ingerlise nørgaard, thomas v. rasmussen, mikkel vognsen, and martin whitehouse provided support and assistance in sample preparation and collection of analytical data. the nordsim laboratory is funded and operated under agreement between the research funding agencies of denmark, norway, and sweden, gtk, finland, and naturhistoriska riksmuseet, sweden; this is nordsim contribution no. 164. clark friend, lotte melchoir larsen and an anonymous reviewer are thanked for critical reviews. references connelly, j.n. & mengel, f.c. 2000: evolution of archean components in the paleoproterozoic nagssugtoqidian orogen, west greenland. geological society of america bulletin 112, 747–763. connelly, j.n., van gool, j.a.m. & mengel, f.c. 2000: temporal evolution of a deeply eroded orogen: the nagssugtoqidian orogen, west greenland. canadian journal of earth sciences 37, 1121–1142. garde, a.a. 2004: geological map of greenland, 1:100 000, kangaatsiaq, 68 v.1 syd. copenhagen: geological survey of denmark and greenland. garde, a.a. in press: geological map of greenland, 1:100 000, ikamiut, 68 v.1 nord. copenhagen: geological survey of denmark and greenland. garde, a.a. & steenfelt, a. 1999: precambrian geology of nuussuaq and the area north-east of disko bugt, west greenland. geology of greenland survey bulletin 181, 6–40. garde, a.a., christiansen, m.s., hollis, j.a., mazur, s. & van gool, j.a.m. 2004: low-pressure metamorphism during archaean crustal growth: a low-strain zone in the northern nagssugtoqidian orogen, west greenland. geological survey of denmark and greenland bulletin 4, 73–76. henderson, g. 1969: the precambrian rocks of the egedesminde–christianshåb area (sheets 68v.1 and 68v.2). rapport grønlands geologiske undersøgelse 23, 1–37. hollis, j.a., garde, a.a., van gool, j.a.m. & thrane k. 2004: polymetamorphism in the northern nagssugtoqidian orogen: a review and presentation of recent data. danmarks og grønlands geologiske undersøgelse rapport 2004/17, 25–27. jacobsen, b.h. 1987: a case for upward continuation as a standard separation filter for potential-field maps. geophysics 52, 1138–1148. kalsbeek, f. & nutman a.p. 1996: anatomy of the early proterozoic nagssugtoqidian orogen, west greenland, explored by reconnaissance shrimp u-pb zircon dating. geology 24, 515–518. kalsbeek, f. & taylor, p.n. 1999: review of isotope data for precambrian rocks from the disko bugt region, west greenland. geology of greenland survey bulletin 181, 41–47. kalsbeek, f., pidgeon, r.t. & taylor, p.n. 1987: nagssugtoqidian mobile belt of west greenland: a cryptic 1850 ma suture between two archaean continents – chemical and isotopic evidence. earth and planetary science letters 85, 365–385. keiding, m. 2004: petrologiske og geokronologiske unders¿gelser af prækambriske bjergarter fra kangilinaaq, vestgrønland, 70 pp. unpublished master thesis, københavns universitet, danmark. kystol, j. & larsen, l.m. 1999: analytical procedures in the rock geochemical laboratory of the geological survey of denmark and greenland. geology of greenland survey bulletin 184, 59–62. ludwig, k.r. 2000: isoplot/ex version 2.2: a geochronological toolkit for microsoft excel. berkeley: berkeley geochronology center. mazur, s., piazolo, s. & alsop, g.i. 2006: structural analysis of the northern nagssugtoqidian orogen, west greenland: an example of complex tectonic patterns in reworked high-grade metamorphic terrains. in: garde, a.a. & kalsbeek, f. (eds): precambrian crustal evolution and cretaceous–palaeogene faulting in west greenland. geological survey of denmark and greenland bulletin 11, 163– 178 (this volume). mottl, m.j. 1983: metabasalts, axial hot springs, and the structure of hydrothermal systems at mid-ocean ridges. geological society of america bulletin 94, 161–180. nicollet, c. & andriambololona, d.r. 1980: distribution of transition elements in crustal metabasic igneous rocks. chemical geology 28, 79–90. nielsen, b.m. & rasmussen, t.m. 2004: mineral resources of the precambrian shield of central west greenland (66° to 70°15′n). part 31 3. implications of potential field data for the tectonic framework. danmarks og grønlands geologiske undersøgelse rapport 2004/ 21, 165 pp. noe-nygaard, a. & ramberg, h. 1961: geological reconnaissance map of the country between latitudes 69°n and 63°45′n, west greenland. meddelelser om grønland 123, 1–9. østergaard, c., garde, a.a., nygaard, j., blomsterberg, j., nielsen, b.m., stendal, h. & thomas, c.w. 2002: the precambrian supracrustal rocks in the naternaq (lersletten) and ikamiut areas, central west greenland. geology of greenland survey bulletin 191, 24– 32. pearce, n.j.g., perkins, w.t., westgate, j.a., gorton, m.p., jackson, s.e., neal., c.r. & cheney, s.p. 1997: a compilation of new and published major and trace element data for nist srm 610 and nist srm 612 glass reference material. geostandards newsletter 21, 115–144. piazolo, s., alsop, g.i., nielsen, b.m. & van gool, j.a.m., 2004: the application of gis to unravel patterns of deformation in high grade terrains: a case study of indentor tectonics from west greenland. in alsop, g.i. & holdsworth, r.e. (eds): flow processes in faults and shear zones. geological society special publication (london) 224, 63–78. shervais, j.w. 1982: ti-v plots and the petrogenesis of modern and ophiolitic lavas. earth and planetary science letters 59, 101–118. taylor, p.n. & kalsbeek, f. 1990: dating the metamorphism of precambrian marbles: examples from proterozoic mobile belts in greenland. chemical geology 86, 21–28. taylor, s.r. & mclennan, s.m. 1985: the continental crust: its composition and evolution, 312 pp. oxford: blackwell scientific publications. thorning, l. 1993: project aeromag-92: a new high resolution aeromagnetic survey of the lersletten area, central west greenland (68°15′ to 68°55′n, 50°25′ to 53°35′w). open file series grønlands geologiske undersøgelse 93/2, 36 pp. thrane, k. & connelly, j.n. 2006: zircon geochronology from the kangaatsiaq–qasigiannguit region, the northern part of the 1.9– 1.8 ga nagssugtoqidian orogen, west greenland. in: garde, a.a. & kalsbeek, f. (eds): precambrian crustal evolution and cretaceous– palaeogene faulting in west greenland. geological survey of denmark and greenland bulletin 11, 87–99 (this volume). van gool, j.a.m. 2005: geological map of greenland, 1:100 000, kangersuneq, 68 v.2 syd. copenhagen: geological survey of denmark and greenland. van gool, j.a.m., connelly, j.n., marker, m. & mengel, f.c. 2002a: the nagssugtoqidian orogen of west greenland: tectonic evolution and regional correlations from a west greenland perspective. canadian journal of earth sciences 39, 665–686. van gool, j.a.m. et al. 2002b: precambrian geology of the northern nagssugtoqidian orogen, west greenland: mapping in the kangaatsiaq area. geology of greenland survey bulletin 191, 13–23. whitehouse, m.j., claesson, s., sunde, t. & vestin, j. 1997: ion microprobe u-pb zircon geochronology and correlation of archaean gneisses from the lewisian complex of gruinard bay, northwestern scotland. geochimica et cosmochimica acta 61/20, 4429–4438. wiedenbeck, m., alle, p., corfu, f., griffin, w.l., meier, m., oberli, f., von quadt, a., roddick, j.c. & spiegel, w. 1995: three natural zircon standards for u-th-pb, lu-hf, trace element and ree analyses. geostandards newsletter 19/1, 1–23. __________________________________________________________________________________________________________________________________________________________________________________ manuscript received 4 october 2004; revision accepted 15 february 2006 32 geological survey of denmark and greenland bulletin 17, 2009,57-60 57 geological maps are of vital importance for documenting and advancing geological knowledge and they are a prerequisite for any meaningful evaluation of economic resources. in greenland, mapping is taking place on the mainland – that for two centuries has been the traditional exploration target – and offshore, where only in the last decades has hydrocarbon exploration moved to the continental shelves. greenland with its 2 166 000 km2 is the largest island in the world. however, the land is overwhelmed by ice. a central ice sheet – the inland ice – blankets some 81% of the country reducing rock outcrop to a coastal fringe 0 to 300 km wide (fig. 1). the continental shelves comprise a little more than twice the area of this fringe, c. 830 000 km2. this preamble serves to emphasise that greenland’s three physiographic units – exposed fringe, offshore and inland ice – are of very different size and that mapping has focused on the smallest acreage. piecing together the composition of the largest, and hitherto unexplored, unit constitutes the next chapter of greenland mapping. historical perspective and aim of this paper in the last 25 years, great strides have been made in geolo gical understanding as can be seen from two 1:2 500 000 maps (escher 1970; escher & pulvertaft 1995). apart from the progress recorded in the ice-free fringe, the 1995 map provides a first interpretation of the offshore, and it also includes information of sub-ice bedrock although this is but a single blob of colour at borehole gisp 2 (fig. 1). this paper’s aim is to provide a first graphic interpretation of the bedrock under the inland ice and to review data sources. its four-page limit does not allow citation of specific sources; these will be covered in a forthcoming paper. this state-of-the-art map is naturally rudimentary in approach with all boundaries arbitrary but it has the prospect of directing attention to future data assembly. the state of knowledge 2008 present knowledge of sub-ice geology is based on six main sources, each discussed below with emphasis on its use in compilation of the geological map shown in fig. 1. drill sites drilling through the ice is the ultimate way of determining substratum composition. however, the only in situ rock sampled is from borehole gisp 2 – an archaean granitoid rock reactivated during the palaeoproterozoic (fig. 1). other ice cores have revealed information about rock debris, for example, camp century (fountain et al. 1981). nunataks nunataks are restricted to the inland ice margin within c. 30 km of the nearest land. most expose locally known rocks and are important for piecing together structural make-up. of importance for the new map is the 120 km long n–s string of nunataks west of dronning louise land, north-east greenland. they infer larger sub-ice occurrences of meso proterozoic sediments than exist on the neighbouring land. coast to coast correlation greenland’s tapering form enables precambrian rocks to be correlated across its southern tip and, by extrapolation farther north, under the ice. the presence on both coasts of archae an rocks flanked north and south by palaeo pro ter o zoic orogenic belts allows the sub-ice projection of the north atlantic craton although its southern and northern boundaries are hidden for 250 km and 500 km, respectively. this correlation is strengthened by aeromagnetic data (fig. 2). the disappearance of palaeogene extrusives on both sides of the inland ice might suggest a single province. however, both coasts are eruption sites connected to continental breakup but since the role of plumes and hot-spots is still unclear, basalts cannot be excluded from central greenland. glacial erratics the inland ice is a relic of a vast pleistocene ice cover and the surrounding land is strewn with rocks dropped as the ice retreated. broadly speaking, erratic suites from southern greenland represent extensively exposed precambrian and late phanerozoic provinces whereas farther north, exotic precambrian–palaeozoic suites relate to sub-ice occurrences that are unknown or not exposed locally. moreover, even the © geus, 2009. geological survey of denmark and greenland bulletin 17, 57–60. available at: www.geus.dk/publications/bull the bedrock geology under the inland ice: the next major challenge for greenland mapping peter r. dawes rosa_2008:rosa-2008 01/07/09 15:48 side 57 58 r i n k i a n c a l e d o n i a n camp century washington land gisp 2 e l l e s m e r e – i n g l e f i e l d c o m m i t t e e – m e l v i l l e peary land dronning louise land v i c t o r i a e l l e s m e r i a n north atlantic craton n a g s s u g to q i d i a n nuuk melv i l l e bugt tasiusaq victoria fjord k e t i l i d i a n ? ? ? nares strait f r a n k l i n i a n baffin island labrador grenville orogen ellesmere island c a n a d a g r e e n l a n d labrador sea baffin bay mesoproterozoic intrusions precambrian shield ellesmerian (devonian) caledonian (silurian) phanerozoic basins devonian–palaeogene cambrian–silurian proterozoic basins with basalts mesoproterozoic– ?neoproterozoic mesoproterozoic, porphyries, redbeds ?mesoproterozoic, redbeds gardar palaeo–neoproterozoic sub-ice basins basalts and intrusions phanerozoic fold belts palaeogene volcanic province palaeoproterozoic palaeoproterozoic, with reworked archaean mainly archaean (palaeoproterozoic overprint) archaean incl. palaeoarchaean (palaeoproterozoic overprint) meso-neoarchaean archaean including eoarchaean thrust deep borehole 250 km kronprins frederik land kn ud r asm uss en land margin of inland ice fig. 1. geological map of greenland with interpretation of sub-ice bedrock in terms of major provinces. ice-free geology (in dark colour shades) modified from henriksen (2008); dashed, grey line, division of proterozoic crust from dahl-jensen et al. (2003). small map, canadian– greenland correlations in the precambrian shield showing the palaeozoic franklinian basin blanketing its northern margin. rosa_2008:rosa-2008 01/07/09 15:48 side 58 59 absence of particular erratics can be informative, for example, this author has no knowledge of erratics that might indicate a late palaeozoic – mesozoic sub-ice source. information from five erratic suites is incorporated into fig. 1. glacial drift across the proterozoic–phanerozoic platform of north greenland is characterised by shield blocks – granitoid rocks, gneisses and associated rocks. most of these need laboratory work to cast light on their age and use in reflecting hidden provinces (see 1 below) but some rocks are ready-made indicators (2, 3). farther south, exotic suites occur on the shield terrain of the west and east coasts (4, 5). 1. the shield erratics isotopically dated are from peary land and environs and they suggest sub-ice archaean crust affected by strong palaeoproterozoic overprint. 2. banded iron formation (bif) characterises the neo archaean committee–melville orogen of baffin is land and north-west greenland. bif erratics in north green land suggest an extension of this terrane far to the east. 3. erratics of porphyries and basalt, with rare sandstone, in washington land indicate a sub-ice volcanic–redbed province (fig. 3a). preliminary isotopic work points to a meso proterozoic age. 4. red sandstone and siltstone erratics around tasiusaq and farther north in north-west greenland point to extensive sub-ice sources (fig. 3b). 5. erratics along the east greenland ice margin were emphasised by haller (1971, fig. 48): proterozoic sandstone and basalt, cambrian quartzite with skolithos and ordovician limestone infer extensive sub-ice sources. detrital provenance studies age and palaeoflow history of detritus within sedimentary rocks – rock clasts and crystals – can be relevant for sub-ice geology. however, minerals like zircon can be transported thousands of kilometres before deposition and identifying sub-ice geology on grains alone is problematical. thus, the clast–grain couplet of the tilloidal neoproterozoic morænesø formation in southern peary land is relevant, particularly so with its south-westerly provenance (kirkland et al. 2009). clasts are of local mesoproterozoic sandstone and dolerite, with less frequent granitoid rocks, bif and porphyry, suggesting proximal sub-ice sources of neoarchaean and meso proterozoic ages. age estimates of zircons from granitoid clasts are 2.7 ga with overprinting at 1.25 ga. zircon crystals range from palaeo archaean to mesopro terozoic with strong palaeo proterozoic peaks suggesting large sub-ice areas. a minor 3.3 ga peak is an obvious link to the substratum (victoria fjord complex) that contains the only known rocks of this age in greenland (nutman et al. 2008). moreover, the subsidiary status of these grains compared to neoarchaean also exists in sequential mesoproterozoic and cambrian strata implying that palaeo archaean rocks form but a minor com ponent of the complex. kirkland et al. (2009) favour two south-western sources for meso proterozic detritus: proximal sub-ice gren ville-overprinted rocks and the type grenvillian of labrador, more than 2000 km distant. a third source is suggested by fig. 1: the sub-ice volcanic province that may also source the rare porphyry clasts. geophysics geophysical methods – satellite, airborne or ice based – undoubtedly have great potential for mapping the sub-ice geology. preliminary interpretations about structure and fig. 2. grids of total magnetic field over southern greenland and offshore based on verhoef et al. (1996, low resolution maanaoala data, geol ogical survey of canada) and white frame, high resolution aeromag data (geological survey of denmark and greenland). stippled lines, geol ogical trends by b.m. stensgaard (personal communication 2009). fig. 3. glacial erratics from sub-ice provinces unknown in outcrop. a: feldspar porphyry from washington land, western north greenland, ggu 425204. other porphyry erratics are illustrated in dawes et al. 2000, fig. 3. b: coarse-grained, cross-bedded sandstone from tasiusaq area, north-west greenland, ggu 457508. photos: jakob lautrup. 553 289 181 103 50 10 -23 -54 -86 -119 -158 -206 -326 363 174 111 70 42 16 -7 -30 -52 -77 -105 -141 -210 a b rosa_2008:rosa-2008 01/07/09 15:48 side 59 crustal thickness can be made from regional magnetic and gravity surveys. the power of aeromagnetics is illustrated by verhoef et al.’s (1996) compilation of reconnaissance data that shows arcuate coast to coast anomalies coinciding with tectonic segments of the nagssugtoqidian orogen, while mergence with high-resolution coastal data allows some structural subdivision of the shield (fig. 2). radar and remote-sensing techniques provide physiographic details about the sub-ice landscape, for example, legarsky et al.’s (1998) work used to locate the volcanic province in knud rasmussen land (fig. 1). many tectonic provinces display distinct physiographical characteristics and thus 3-d imagery is vital for mapping sub-ice geology. mountains, plains, plateaux and lowlands are not the only geological indicators, but hills and valleys affect ice dynamics and control water flow, two primary parameters for determining provenances of erratics and detrital material. conclusions, future research, ice recession and economic potential a main conclusion must be that while mapping below the ice is in its infancy, the status of the geus databases has promising potential for planning research, whether sampling, drill ing or geophysics. one dire need is for low-altitude and ice-based geophysical surveys to facilitate deductions about spatial relationships of sub-ice provinces. the new map leads to eight conclusions, but being conjectural, it raises important questions – too many to discuss in this short paper. 1. provinces unknown in outcrop occur below the ice. 2. where it is widest, in the north, the inland ice hides the most variable geology: palaeoarchaean, neoarchaean, pala eoproterozoic, mesoproterozoic and palaeozoic pro vinces. 3. archaean crust underlies kronprins frederik land but its eastern connection is unknown. 4. whether the ellesmere–inglefield juvenile crust links genetically and structurally (or at all) with palaeo pro ter ozoic rocks within the caledonian fold belt remains open. 5. the volcanic province of knud rasmussen land reinforces the profusion of proterozoic rift-related magmatism along the rim of the north american craton. 6. mesoproterozoic–ordovician rocks are widespread be yond the caledonian front linking northern foreland outcrops to sub-ice occurrences in central east greenland. 7. the potential for sub-ice basins of late palaeozoic – mesozoic age is limited. 8. the presence of palaeogene volcanic rocks in central greenland cannot be dismissed. currently, greenland plays centre stage in the global climate debate, its recessive ice margin with spectacular, shrinking glaciers being international attractions. with this coveted popularity come startling prophecies, for example, “as its huge ice sheets begin to melt, it [greenland] could find itself sitting on a fortune in oil and gems” (barkham 2008). be this as it may, before the rocks of the hidden 81% have been mapped, assessments of economic potential – often judged poor compared with neighbouring canada despite common geology (fig. 1) – remains equivocal. references barkham, p. 2008: beyond the ice. the guardian, 11 december 2008, electronic version. london: guardian news & media. dahl-jensen, t., larsen, t.b., woelbern, i., bach, t., hanka, w., kind, r., gregersen, s., mosegaard, k., voss, p. & gudmundsson, o. 2003: depth to moho in greenland: receiver-function analysis suggests two proterozoic blocks in greenland. earth and planetary science letters 205, 379–393. dawes, p.r., thomassen, b. & andersson, t.i. 2000: a new volcanic province: evidence from glacial erratics in western north greenland. geology of greenland survey bulletin 186, 35–41. escher, a. 1970: geological/tectonic map of greenland, 1:2 500 000. copenhagen: geological survey of greenland. escher, j.c. & pulvertaft, t.c.r. 1995: geological map of greenland, 1:2 500 000. copenhagen: geological survey of greenland. fountain, j., usselman, t.m., wooden, j. & langway, c.c. 1981: evidence of the bedrock beneath the greenland ice sheet, near camp century, greenland. journal of glaciology 27(95), 193–197. haller, j. 1971: geology of the east greenland caledonides, 413 pp. new york: interscience publishers. henriksen, n. 2008: geological history of greenland, 272 pp. copen hagen: geological survey of denmark and greenland. kirkland, c.l., pease, v., whitehouse, m.j. & ineson, j.r. 2009: pro venance record from mesoproterozoic–cambrian sediments of peary land, north greenland; implications for the ice-covered shield and laurentian palaeogeography. precambrian research 170, 43–60. legarsky, j., wong, a., akins, t. & gogineni, s.p. 1998: detection of hills from radar in central-northern greenland. journal of glaciology 44(146), 182–184. nutman, a.p., dawes, p.r., kalsbeek, f. & hamilton, m.a. 2008: palaeoproterozoic and archaean gneiss complexes in northern greenland: palaeoproterozoic terrane assembly in the high arctic. precambrian research 161, 419–451. verhoef, j., macnab, r., roest, w.r. & arjani-hamed, j. 1996: magnetic anomalies of the arctic and north atlantic oceans and adjacent land areas. geological survey of canada, open file report 3125a, 225 pp. author’s address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: prd@geus.dk 60 rosa_2008:rosa-2008 01/07/09 15:48 side 60 geological survey of denmark and greenland bulletin 38, 2017, 25-28 25 danish glacial landscape elements such as basal till plains, hummocky moraine areas and outwash plains contain a variety of small and large depressions. they were probably formed in glacial, late-glacial or holocene time and may represent dead-ice holes or degraded pingos, or sinkholes formed by interaction between pre-quaternary chalk or limestone bedrock and the thin glacial co ver. the aim of this study is to map terrain depressions that might potentially be karst sinkholes by analysing digital terrain models in the geographic information system (gis). the incentive to apply the technique for mapping of sinkholes came from an accidental acquaintance with a farmer, jens kirk, whose farmland is located near thisted. jens kirk told us that the front end of his tractor had suddenly sunk into the ground during routine farming work, and this incident was our inspiration to start the project described here. geological setting and study areas sinkholes have previously been registered in mainly northern denmark where the pre-quaternary surface consists of chalk and limestone (feddersen 1880). in these areas sinkholes and other karst features including dolines, karst lakes, small caves, disappearing streams and karst springs are known (nilsson & gravesen 2017). the landforms are found at locations where quaternary surface layers are thin and calcareous rocks are present near the surface. sinkholes can be formed when acid groundwater dissolves and removes calcium carbonate from the calcareous sediment. cavities are formed in the carbonate sediments, and the overlying deposits may collapse and create a surface depression. the present authors have carried out pilot studies around limfjorden and in the northern part of djursland in jylland. the results from north of thisted and in the planted wood of svinkløv plantage south of svinkløv are presented here (fig. 1). it should be noted, that svinkløv plantage has never been farmed, and the depressions found here are small compared to those in farmed areas. this means that even small depressions have been preserved if not covered by drifting sand. the first study area is located 2–3 km north of thisted (fig. 1). the pre-quaternary surface of maastrichtian chalk and danian bryozoan limestone and micritic limestone in this area is commonly covered by a thin (less than 5 m thick) quaternary cover. small dolines filled with paleocene clay are found in the danian limestone in karst sinkhole mapping using gis and digital terrain models peter brøgger sørensen, holger lykke-andersen, peter gravesen and bertel nilsson •••••••••••••••••••••••••••••••••••••••• ••••••••••••••••••• ••• ••••••• ••••••••••••••••••••••••••• ••••••••••• ••••••••••••••• •••••••••••••••••••••• •• • •••••••••••••• •• •••••• ••• • ••••••••••••• •••••••• •••• • •••••• ••• •••••• ••• •••• •• ••• ••••• ••••••••••••••• ••••• ••••••••• •• • •••••••••••••••••••••••••••••••••••• • •••• •• •• ••••••• •••••••• •••• • ••• ••• •• ••••••••••••••• •• ••••••••••••••••••••••••••••••••••••••••••••••••••••••••••••••••• •• 40 km pre-quaternary deposits møns klint fm (upper cretaceous) mapped sinkholes sinkholes: dhm 0.4 m data sinkholes: dhm 1.6 m data miocene oligocene eocene paleocene københavns kalk fm (danian) stevns klint fm (danian) lower cretaceous and older svinkløv plantage djursland jylland skive lim fjor den thistedthisted fjerritslevfjerritslevnors sø 100 km denmarkdenmark sweden fig. 1. locations of mapped small depressions in northern jylland. the surface is shown as pre-quaternary where parts of it consists of cretaceous chalk and danian limestone covered by quaternary deposits. red and blue dots: possible sinkholes mapped with the 1.6 and 0.4 m digital elevation models, respectively. black frames: local study areas. © 2017 geus. geological survey of denmark and greenland bulletin 38, 25–28. open access: www.geus.dk/publications/bull 2626 thisted town (gravesen & jakobsen 2016). sinkholes are distributed over the entire area. in cliffs and limestone pits it is possible to find two types of dolines: funnel-shaped and vertical-sided. both types have originated as areas of intense fracturing where acid water has percolated downwards and dissolved the chalk. in the western part of the area, large karstic lakes such as nors sø occur along a fault zones and in the middle of a salt structure (hansen & håkansson 1980). between svinkløv and fjerritslev on the north-west coast of jylland many sinkholes and dolines are registered in the chalk and limestone cliffs (gry 1979). sinkholes also occur in the area of elevated chalk south of fjerritslev. methods and data the detailed topographic maps that became available from c. 1900 made it possible for ussing (1903) and harder (1908) to interpret the danish glacial landscape. since then the technological evolution has resulted in a highly detailed digital terrain model (in danish ‘digital højdemodel’, dhm) with a lateral resolution of 1.6 m in the 2007 version (dhm 1.6) and 0.40 m in the 2016 version (dhm 0.4) with a vertical accuracy of a few centimetres (www.sdfe.dk). using hill-shade versions of the dhm 1.6 model with a vertical exaggeration of 3 even the smallest vertical variations can be detected. numerous depressions were identified, and detailed mapping was initiated, using qgis (open source version of gis) as a data repository. the dhm 0.4 model is delivered with hill shade but with no vertical exaggeration. the 3d geographical information of the mapped depressions was analysed in gis supported by other sources of spatial data such as orthophoto and topographic maps. mobile devices such as tablets and smartphones with gps were used in the field documentation of the depressions throughout the mapping project. two of the authors recently demonstrated the usefulness of the combination of these techniques for identification and mapping of remnants of pingos in denmark (sørensen & lykke-andersen 2017). more than 350 depressions supposed to be sinkholes in northern jylland were mapped (fig. 1); in svinkløv plantage sinkholes were mapped using the dhm 0.4 model. the digital terrain models were provided by the danish agency for data supply and efficiency (www.sdfe.dk) as a digital service in raster format with and without hill shade for visual interpretation, and as a web map service for analysis and measurement. the depression mapping was divided into three steps. firstly, the location of each depression was found and digitised using gis. secondly, a profile of the depression was constructed using the qgis profile tool and its diameter and depth measured, allowing a descriptive statistical analysis of the sinkholes and their spatial distribution. inspection of orthophotos was sometimes helpful because the vegetation pattern can reveal differences in the topsoil. on this basis the depression was either approved or rejected. for example, a shallow depression on an inclined surface facing away from the incident light produces a significant shadow which does not represent a sinkhole. finally, a field check was carried out for some of the proposed sinkholes, because previous studies have shown that the structural and sedimentary conditions of some of the depressions should be investigated by excavation (zhu et al. 2014). 500 m 500 m a b fig. 2. hill-shade map with 10 × vertical exaggeration showing examples of sinkholes mapped with the 1.6 m digital elevation model. a: unmarked map. b: mapped sinkholes (red dots). yellow dot: location of the excavated ‘jens kirk sinkhole’ north of thisted as shown in fig. 3. 27 results results from the depression mapping north of thisted and in svinkløv plantage are shown in figs 2 and 3. in the thisted area c. 116 sinkholes are identified, and 20 depressions were identified on the hill shade map within the subarea measuring 1 × 1.5 km shown in fig. 2. the diameter of the depressions in the thisted area varies from 21 to 100 m, with depths between 0.3 and 2.5 m. to obtain information about the geometry of the depression encountered by jens kirk and to prove it is a sinkhole, a trench was dug (fig. 3). the diameter of the excavated depression is 25 m. outside the depression, the depth below the topsoil to the chalk surface was 30 cm. in the middle of the depression the chalk surface occurs about 2.35 m below the terrain surface outside the sinkhole. in the south-western part of the profile, the chalk surface is almost vertical, indicating subsidence along a fault. stone-free topsoil, probably glacial, was found overlaying the chalk. in the middle of the depression about 40 cm above the chalk surface, a chert layer 7 m long and 5–10 cm thick was found. the geometry and structure of the depression strongly indicate that it is a karstic sinkhole. it should be noted that the diameter of the ‘jens kirk sinkhole’ (fig. 3a) is in the lower range of those mapped north of thisted. in svinkløv plantage, 65 small sinkholes were mapped with success using the dhm 0.4 model (fig. 4), which are not visible in the 1.6 m model; the area is partly covered by dune sand. the diameter of the small sinkholes varies from 1 to 10 m, with depths between 0.4 and 4.5 m. a survey in the bulbjerg area has revealed a few similar small depressions that are likewise supposed to be sinkholes. the dhm 0.4 model is obviously better suited for the minor landscape elements such as the small depressions found in svinkløv plantage even with the absence of vertical exaggeration. a significant difference in the diameter/depth ratio between the sinkholes in the thisted and svinkløv plantage areas is seen in table 1. the sinkholes in svinkløv plan0 4 8 12 16 20 24 26 0 100 200 m cm below surface ne sw glacial sediments maastrichtian chalkcherta b fig. 3. a: section drawn from the excavated trench through the ‘jens kirk sinkhole’ (vertical exaggeration 1.7 ×). b: the excavated trench with the chalk surface exposed at the bottom. utm coordinates: 481294e, 6316209n. 1 km fig. 4. sinkholes (red dots) mapped in svinkløv plantage using the 0.4 m digital elevation model. 2828 tage are ten times smaller but slightly deeper than those in thisted. discussion origin of the mapped depressions mapping of karst sinkholes using gis and dhm in areas with carbonate sediments underlying thin quaternary deposits seems to be evident. however, the mapping in this study also identified depressions in areas with palaeogene clays underneath thin glacial deposits (area west of skive; fig. 1). these depressions cannot have been developed due to karst evolvement. thus, dead-ice holes or degraded pingos can be other potential origins of these depressions and have thoroughly been discussed in sørensen & lykkeandersen (2017). however, it is beyond the scope of this study to collect the required structural and sedimentary documentation to classify all of the individual depressions found in northern jylland. factors influencing the formation and geometry of the mapped sinkholes sinkholes may develop if the pre-quaternary surface consists of carbonate sediments such as maastrichtian chalk or danian limestone, which are easily dissolved and contain fractures, hollows and cavities. in order to allow the sinkhole-forming process to operate during the neogene, the overlying sediments (mainly quaternary deposits) must be thin. the circular shape of the sinkholes is possibly a combination of crossing fault lines and dissolution processes. the anthropogenic activity seems to be an important factor for the collapses, but also the climate with changing wet and dry conditions may influence the development. changes in the groundwater level, abstraction of water and changing drainage patterns are other controlling factors. conclusions 1. the pilot study demonstrates that analysing digital terrain models in gis is a powerful method. the dhm 1.6 model with a vertical exaggeration of 3 is very well suited for the identification of depressions with diameters greater than 20 m, while the dhm 0.4 model is better suited for the identification of small landscape elements even without vertical exaggeration. 2. the sinkholes in svinkløv plantage are much smaller and slightly deeper than in thisted. one explanation for this could be that the svinkløv area has never been farmed. 3. the structural and sedimentary conditions in other sinkholes should be investigated by excavations. references feddersen, a. 1880: nogle danske overfladeforhold. geografisk tidsskrift 4, 113–118. gravesen, p. & jakobsen, p.r. 2016: pre-quaternary rocks and sediments with some of the highest radioactive levels in denmark. bulletin geological survey of denmark and greenland 35, 31–34. gry, h. 1979: description of the geological map of denmark. map sheet løgstør. quaternary deposits. 1:100  000/1:50  000 (in danish with english summary). danmarks geologiske undersøgelse i. række 26, 58 pp. + atlas + map vol. hansen, j.m. & håkansson, e. 1980: thistedstrukturens geologi – et ’neotektonisk’ skoleeksempel. dansk geolologisk forening, årsskrift for 1979, 1–9. harder, p. 1908: en østjydsk israndslinje og dens indflydelse paa vandløbene. danmarks geologiske undersøgelse ii. række 19, 262 pp. nilsson, b. & gravesen, p. 2017: karst geology and regional hydrogeology in denmark. in: white, w.b. et al. (eds): karst groundwater contamination & public health. springer international publishing ag switzerland (in press). sørensen, p.b. & lykke-andersen, h. 2017: kortlægning af pingo-rester i danmark. geologisk tidsskrift 2017, 21–29. ussing, n.v. 1903: om jyllands hedesletter og teorierne for deres dannelse. oversigt over det kongelige danske videnskabernes selskabs forhandlinger 1903(2) 99–152. zhu, j., taylor, t.p., currens, j.c. & crawford, m.m. 2014: improved karst sinkhole mapping in kentuckey using lidar techniques: a pilot study in floyds fork watershed. journal of cave and karst studies 76(3) 207–216. depth, m§ diameter, m§ diameter/depth thisted 0.3–2.5 (0.8) 21–100 (43.6) 57.4 svinkløv plantage 0.4–4.5 1.0–10 (4.3) 3.8 §minimum, maximum and mean values table 1. depths and diameters of sinkholes authors’ addresses p.b.s., møllepold 17, 6200 aabenraa, denmark. e-mail: peter@peterbrogger.dk h.l.-a., department of geoscience, aarhus university, høegh guldbergs gade 2, 8000 aarhus c, denmark. p.g. & b.n., geological survey of denmark and greenland, øster voldgade 10, copenhagen k, denmark. geological survey of denmark and greenland bulletin 13, 2007, 29-32 29 the extensive and very deep ?jurassic/cretaceous–palaeogene sedimentary basins offshore west greenland have a significant petroleum exploration potential. this is particularly true for the offshore region west of disko and nuussuaq where a live petroleum system has been documented for many years. at present, stratigraphic knowledge in this area is almost nonexistent and analogue studies from onshore areas and offshore exploration wells to the south are therefore crucial to understanding the distribution and quality of possible reservoir rocks in the disko–nuussuaq offshore area. one of the main risk parameters in petroleum exploration in this region is the presence of an adequate reservoir rock. tectonostratigraphic considerations suggest that several sand-prone stratigraphic levels are probably present, but their pro v enance and reservoir quality are at present poorly known both onshore and offshore. a sediment provenance study including zircon provenance u-pb dating and wholerock geochemical analysis was therefore initiated by the geological survey of den mark and greenland (geus) in preparation for the disko west licensing round 2006 (scherstén et al. 2007). the main aims of this study were to: 1. characterise the source areas and dispersal patterns for the various sandstone units of cretaceous–paleocene age in the nuussuaq basin and compare these with sandstone units in selected west greenland offshore exploration wells (figs 1, 2), employing advanced zircon provenance u-pb dating using laser ablation inductively coupled plasma mass spectrometry (la-icp-ms; cf. frei et al. 2006). 2. detect possible changes in sediment source with time, e.g. local versus regional sources. provenance of cretaceous and paleocene sandstones in the west greenland basins based on detrital zircon dating anders scherstén and martin sønderholm © geus, 2007. geological survey of denmark and greenland bulletin 13, 29–32. available at: www.geus.dk/publications/bull fig. 1. simplified geological map of eastern canada and greenland (modified from escher & pulvertaft 1995 and st-onge et al. 2006). green land is shown in a paleocene pre-drift position (from oakey 2005, p. 222). arrows indicate possible source of grenvillian age components in west greenland zircon samples. inset shows samp led localities in the nuussuaq basin; 1, itsaku on svartenhuk halvø (sh); 2, upernivik ø; 3, ikorfat; 4, paatuut; 5, kingittoq; 6, atanikerluk; 7, pingu and 8, grønne ejland. sampled wells are gro#3 (g), hellefisk-1 and qulleq-1. 30 zircon as a provenance tool is receiving increasing attention and has proven to be a powerful indicator of clastic sedi ment sources, a tracer of the earth’s oldest materials, and a tracer of continental crust-forming processes (froude et al. 1983; williams & claesson 1987; dodson et al. 1988; fedo et al. 2003; hawkesworth & kemp 2006). zircon is common in continental rocks and it is assumed that its distribution in sediments will normally represent the source rocks. although there are several complications, the sediment zircon u-pb age frequency should in general terms mirror the relative proportions of different source materials. this as sump tion is particularly important if exotic components can be identified, as their frequency will provide an estimate of the exotic influx: it may also be essential in tra cing sediment paths that affect the detrital compositions and subsequent diagenetic history of possible hydrocarbon reservoir rocks. cretaceous sediment provenance it is assumed that the age structure of the north atlantic cratons surrounding the study area is well enough known to constrain the origin of the source components that contributed to the sediments. archaean gneisses that range from 3850 to 2600 ma (hollis et al. 2006) dominate southern west greenland. important peaks occur at 3600, 3100, 2900 and 2700 ma. farther north, the archaean basement was reworked during the nagssugtoqidian/rinkian orogeny, which adds an age peak centred at 1900–1750 ma (figs 1, 3; connelly et al. 2000). as part of this study, 4262 grains were dated from 65 sediment samples from eight localities in the nuussuaq basin and three exploration wells (fig. 1; scherstén et al. 2007). data that are >10% discordant were filtered out as they are more likely to be disturbed by common pb contamination, ancient pb-loss and mixed domains. the remaining 2735 grains display a relative age distribution that is dominated by age peaks between ~2500–3200 ma (fig. 3). there is also a peak at ~3600 ma, which constitutes several samples suggesting that 3600 ma age components are perhaps more abun dant than those from the well-known godthåbsfjord area (friend & nutman 2005; hollis et al. 2006). a ~1900 fig. 2. simplified stratigraphic scheme of the nuussuaq basin and west green land offshore region showing stratigraphic distribution of ana lysed samples. fig. 3. relative 207pb/206pb age distribution of 2735 <10% discordant zircon grains from the cretaceous–palaeogene sediments in this study (red curve). the relative probability reflects the likelihood of finding any given age, although the ~1900 ma peak is overrepresented due to sampling bias. zircon ages from the qulleq-1 well show small but significant peaks at ~1600–1700 and ~1100 ma (n is the number of grains that are within error of 1600 and 1100 ma, respectively). these appear to be coupled and may have been either derived from east greenland or the canadian shield. see text for further discussion. ma peak is also distinct, but is probably overrepresented due to sampling bias through the many samples taken at itsaku where this age component is strongly represented in compa rison to the sediments farther south (see discussion below). the 1900 ma peak is slightly asymmetric with a tail towards younger ages and an age component around 1600–1700 ma. the overall age distribution is in excellent agreement with a source from the west greenland crystalline basement; and the zircon data support the existing depositional models indicating that a major deltaic system drained into the nuussuaq basin from the east-south-east during cretaceous–paleocene time (pedersen & pulvertaft 1992). a small but significant peak at ~1100 ma suggests a distal component that is not readily explained by derivation from the west greenland crystalline basement as described above. this component seems to be associated with the 1600–1700 ma occurrence noted above (fig. 3). two possible sources can explain the dual peaks: from east greenland or labrador (fig. 1). if an east greenland origin is favoured, it would be anticipated that this signature would also be associated with caledonian ages between 380 and 480 ma, which have not yet been identified (figs 1, 3). given the large number of grains analysed, it would be expected that even a very small contribution would have been detected suggesting labrador as the most likely source for the 1100 ma peak. a southern, labrador source for the 1100 ma component in the qulleq-1 well is corroborated by the absence of the 1900 ma peak that is ubiquitous in the nuussuaq basin to the north (fig. 4). the samples from the qulleq-1 well are domi nated by archaean ages without contributions from rocks reworked during the nagssugtoqidian/rinkian orogenic event. in the nuussuaq basin the 1100 ma component is very rare in the onshore, deltaic facies and occurs almost exclusively in the deep-water deposits in accordance with a longshore transport component from the south as the source of this component. however, current data from turbidite channel units on western nuussuaq show transport directions towards the south (dam & sønderholm 1994). it is not possible to explain this apparent discrepancy based on the present database. more data including the other offshore wells will be needed to elucidate the possible interconnections and transport paths in the west greenland offshore basins. paleocene point source provenance on svartenhuk halvø a set of samples was collected on itsaku on svartenhuk halvø where a major hiatus separates an upper albian to lower ceno manian deltaic succession from an upper campanian to paleo cene marine turbidite succession. the detrital zircon age distribution in the deltaic succession is typical for the nuussuaq basin and dominated by a distinct peak at ~2800 ma; this is flanked by scattered peaks between 2400 and 3200 ma (fig. 4), as well as significant peaks at 3600–3800 ma representing eoarchaean components. a 1900 ma peak is another typical feature of the nuussuaq basin sedi ments, whereas the occurrences of 1100 and 1600–1700 ma peaks are more intermittent (see above). the overall pattern is in good agreement with that of the general west greenland cretaceous population and deltaic deposition from the eastsouth-east. the zircon age distribution of the overlying turbidite succession is in stark contrast to the lower section (fig. 4). here, the zircon population forms a single, well-defined ~1900 ma peak with an apparent normal distribution indicative of a single point source with respect to zircon. the only known source that fits this distribution is the 1869 ± 9 ma prøven igneous complex (fig. 1; thrane et al. 2005). assuming the prøven igneous complex forms a single source to the upper part of the succession, a tukey’s biweight mean of 1872 ± 4 ma (n = 275) can be calculated, which is in excellent agreement with the prøven igneous complex. a few grains scatter towards 2700 ma, which likely reflects inherited components, in accordance with its derivation from lower continental crust (thrane et al. 2005). trace element systematics in zircon may provide further constraints on zircon origin in provenance studies (hoskin & ireland 2000). however, several hurdles need to be overcome. for instance, many features are shared intimately between zircon that are derived from widely different sources (hoskin & ireland 2000), and in a detrital population each grain has 31 fig. 4. relative probability diagram for 207pb/206pb ages for zircon from upper albian – cenomanian deltaic sediments (red) and upper cam panian/maastrichtian – paleocene marine sediments (blue) on itsaku, svartenhuk halvø. the deltaic deposits display a pattern that is typical for the sediments in the nuussuaq basin, whereas the constrained pattern of the overlying marine deposits is unique and seems to require a single point source. 32 to be treated separately, as a common age does not necessarily imply the same source rock, and by inference, the same cry stallisation processes or conditions. nevertheless, it is as sumed here that zircon from the upper part of the section represents one population that was derived from the prøven igneous complex. trace element abundances were determined for 138 zircon grains from this part of the section and contrasted against archaean (n = 424) zircon grains from other stratigraphic levels. as expected, there are no systematic variations with age, which would require an age dependent systematic change in zircon crystallisation processes. tita n ium in zircon thermometry (watson et al. 2006) enables calculation of zircon crystallisation temperatures. the upper zircon population yields a mean temperature of 845 ± 19°c (± 2σmean; n = 93; gj-1 standard reproducibility 668 ± 30°c 2σ n = 17), which is similar to zircon saturation temperatures of 790– 880°c (n = 4) calculated from bulk rock data and the independently estimated intrusion temperature of the prøven igneous complex (thrane et al. 2005). the prøven igneous complex is inferred to have been derived from a lower continental crust source as reflected by significantly negative euanomalies from a plagioclase residue (eu/eu* ~0.45; thrane et al. 2005). this appears to be reflected by the zircon population, which has a mean of 0.23 ± 0.03 (± 2σmean; n = 93; gj-1 standard reproducibility 0.96 ± 0.06 2σ n = 23). the values contrast with the average archaean population (eu/eu* = 0.49 ± 0.03 2σmean; n = 257), which is assumed to be dominated by rocks such as tonalite–trondhjemite–grano diorite (ttg) suites that have eu/eu* ~1.0. thus, the zircon population of the upper campanian to paleocene marine turbidite succession seems to form a single population that is in accordance with derivation from the prøven igneous complex. this implies a major change in depositional transport direction compared to the underlying lower cretaceous deltaic deposits from a south-eastern and eastern source to a northern source. acknowledgement the project was supported by the bureau of minerals and petroleum, government of greenland. references connelly, j.n., van gool, j.a.m. & mengel, f.c. 2000: temporal evolution of a deeply eroded orogen: the nagssugtoqidian orogen, west greenland. canadian journal of earth sciences 37, 1121–1142. dam, g. & sønderholm, m. 1994: lowstand slope channels of the itilli succession (maastrichtian – lower paleocene), nuussuaq, west green land. sedimentary geology 94, 49–71. dodson, m.h., compston, w., williams, i.s. & wilson, j.f. 1988: a search for ancient detrital zircons in zimbabwean sediments. journal of the geological society (london) 145, 977–983. escher, j. & pulvertaft, t.c.r. 1995: geological map of greenland, 1 : 2 500 000. copenhagen: geological survey of greenland. fedo, c.m., sircombe, k.n. & rainbird, r.h. 2003: detrital zircon analysis of the sedimentary record. in: hanchar, j.m. & hoskin, p.w.o. (eds): zircon. reviews in mineralogy and geochemistry 53, 277–303. frei d., hollis j.a., gerdes a., harlov d., karlsson c., vasquez p., franz g., johansson l. & knudsen, c. 2006: advanced in-situ trace element and geochronological microanalysis of geomaterials by laser ablation techniques. geological survey of denmark and greenland bulletin 10, 25–28. friend, c. & nutman, a. 2005: new pieces to the archaean terrane jigsaw puzzle in the nuuk region, southern west greenland: steps in transforming a simple insight into a complex regional tectonothermal model. journal of the geological society (london) 162, 147–162. froude, d.o., ireland, t.r., kinny, p.d., williams, i.s., compston, w., williams, i.r. & myers, j.s. 1983: ion microprobe identification of 4000–4200 myr-old terrestrial zircons. nature 304, 616–618. hawkesworth, c.j. & kemp, a.i.s. 2006: evolution of the continental crust. nature 443, 811–817. hollis, j.a., frei, d., van gool, j.a.m., garde, a.a. & persson, m. 2006: using zircon geochronology to resolve the archaean geology of southern west greenland. geological survey of denmark and greenland bulletin 10, 49–52. hoskin, p.w.o. & ireland, t.r. 2000: rare earth element chemistry of zircon and its use as a provenance indicator. geology 28, 627–630. oakey, g.n. 2005: cenozoic evolution and lithosphere dynamics of the baffin bay – nares strait region of arctic canada and greenland, 233 pp. amsterdam: vrije universiteit. pedersen, g.k. & pulvertaft, t.c.r. 1992: the nonmarine cretaceous of the west greenland basin, onshore west greenland. cretaceous research 13, 263–272. scherstén, a., sønderholm, m. & steenfelt, a. 2007: provenance of west greenland cretaceous and paleocene sandstones and stream sediment samples based on u-pb dating of detrital zircon: data and results. dan marks og grønlands geologiske undersøgelse rapport 2007/21, 121 pp. st-onge, m. (compiler), van gool, j.a.m., garde, a.a. & scott, d.j. 2006: correlations of archaean to mesoproterozoic units and structures across baffin bay, davis strait and the labrador sea: kangerlussuaq workshop 2005 report and literature review. danmarks og grønlands geologiske undersøgelse rapport 2006/6, 20–57. thrane, k., baker, j., connelly, j. & nutman, a. 2005: age, petrogenesis and metamorphism of the syn-collisional prøven igneous complex, west greenland. contributions to mineralogy and petrology 149, 541–555. watson, e.b., wark, d.a. & thomas, j.b. 2006: crystallisation thermometers for zircon and rutile. contributions to mineralogy and petrology 151, 413–433. williams, i.s. & claesson, s. 1987: isotopic evidence for the precambrian provenance and caledonian metamorphism of high grade paragneisses from the seve nappes, scandinavian caledonides: ii. ion microprobe zircon u-th-pb. contributions to mineralogy and petro logy 97, 205–217. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: asch@geus.dk geological survey of denmark and greenland bulletin 17, 2009, 21-24 the triassic – lower cretaceous sedimentary succession of the norwegian–danish basin has for a long time been of exploration interest, and numerous studies have been carried out. however, high-resolution correlation within the basin remains necessary, especially between the danish and nor we gian parts of the basin. a variety of lithoand biostratigraphic schemes have been applied to the succession over the years, but lack of consistency in terminology has often led to confusing interpretations of the geological development. in this study a sequence stratigraphic scheme has been developed for the danish basin and a compiled palynological event stratigraphy is applied to a number of wells connecting the danish and norwegian parts of the basin and new marker horizons are identified. one of the aims of this study is to reach consistency in order to facilitate correlation within the basin and we also emphasise the recognition of a potentially important mid-triassic event in the basin. geological setting the intracratonic permian–cenozoic norwegian–danish basin is bounded to the south by the ringkøbing–fyn high and to the north by the strongly faulted sorgenfrei–tornquist zone (fig. 1). the basin was formed by late carboniferous – early permian crustal extension followed by thermal sagging, local faulting and salt tectonics. the syn-rift succession consists of rotliegendes volcaniclastic rocks, alluvial conglomerates and sandstones as well as lacustrine mudstones. the overlying zechstein–cenozoic post-rift succession consists of two major sequences separated by an early middle jurassic unconformity that reflects regional uplift and erosion (nielsen 2003). the lower sequence comprises zechstein salt, triassic sandstones, mudstones, marls and carbonates and lower jurassic claystones, while the upper sequence encompasses middle jurassic – lower cretaceous clastic rocks, ladinian palynofloras in the norwegian–danish basin: a regional marker reflecting a climate change sofie lindström, henrik vosgerau, stefan piasecki, lars henrik nielsen, karen dybkjær and mikael erlström © geus, 2009. geological survey of denmark and greenland bulletin 17, 21–24. available at: www.geus.dk/publications/bull 21 100 km land redbeds sea fault a-2 höllviken-2 ffc-1/2 8/11-1 felicia-1 100 m 8/11-1 west 9/11-1 10/7-1 11/10-1 inez-1 f-1x east k-1x felicia-1 børglum fm flyvbjerg fm haldager fm fj er ri ts le v fm gassum fm vinding fm oddesund fm skagerrak fm f-iii f-iv f-ii f-i gr sonic gr sonicgr sonic gr sonic gr sonic gr sonic gr sonic gr sonicflyvbjerg fm gassum fm marine mudstones shallow marine sandstones and siltstones, offshore mudstones paralic and non-marine sandstones, mudstones and coals alluvial conglomerates, sandstones and lacustrine mudstones marine calcareous claystones, carbonates ladinian palynofloras sabkhas and lacustrine calcareous, evaporitic mudstones danish wellsnorwegian wells fyn high ringkøbing – norwegian–danish basin sorgenfrei–tornquist zone fig. 1. correlation of middle triassic to upper jurassic strata along a transect. the map shows the location of the wells and the middle triassic palaeogeography. gr, gamma ray log; sonic, sonic log. rosa_2008:rosa-2008 01/07/09 15:48 side 21 upper cretaceous chalk and cenozoic clastic rocks. large parts of the lower sequence are difficult to date accurately by biostratigraphy owing to the predominantly continental strata deposited during a hot and arid climate, whereas the upper sequence is dominated by marine fossiliferous de po sits, which are easier to date. the investigated succession and the stratigraphic approach the principal reservoir rocks and potential source rocks of the basin were formed during late triassic – late jurassic times (petersen et al. 2008). different lithostratigraphic terminologies are used for the norwegian and danish areas (fig. 2). an inconsistent mixture of lithostratigraphic names from both norwegian and danish schemes is commonly used by the operators in well reports, and in some cases, lithostratigraphic names normally applied to units in the norwegian central graben further contribute to the confusion. a robust sequence stratigraphic scheme established in the danish basin by nielsen (2003) and a compiled palynostratigraphic event scheme are applied to a number of wells (felicia-1, k-1, f-1, inez-1, 11/10-1, 10/7-1, 9/11-1 and 8/11-1) situated along an e–w trending transect from the fjerritslev fault nw of jylland to the border of the norwegian central graben (fig. 1). based on characteristic well-log patterns supported by lithological descriptions from well reports, formation tops and sequence stratigraphic surfaces were identified and integrated with palynological events identified in this study. the latter events are based on established spore–pollen and dinoflagellate cyst stratigraphies for the triassic – early cretaceous of western europe. selected im portant palynostratigraphic events and sequence stratigraphic surfaces are shown in fig. 2. 22 oddesund fm se ri es sy st em stage ju ra ss ic t ri as si c norwegian–danish basin børglum formation flyvbjerg formation haldager sand formation gassum formation vinding formation skagerrak formation muschelkalk fm frederikshavn formationvolgian kimmeridgian oxfordian callovian bathonian bajocian aalenian toarcian pliensbachian sinemurian hettangian rhaetian norian carnian ladinian fjerritslev fm sauda fm tau fm egersund fm sandness fm bryne fm u pp er m id dl e m id dl e lo w er u pp er mfs 15 sb 19 sb 20 sb 21 sb 22 sb 23 sb 24 sb 15 sb 11 mfs 7 mfs 1 sb 5 ts 1 ts 22 ts 23 sb 9 paralic and non-marine sandstones, siltstones, mudstones and coals marine mudstones and siltstones unconformity shallow marine sandstones and siltstones, offshore mudstones alluvial sandstones and lacustrine mudstone marine calcareous claystones, carbonates hiatus sequence boundary maximum flooding surface transgressive surface danish part norwegian part selected key surfaces last occurrence datums gassum fm skagerrak formation smith bank formation mfs ts sb fjerritslev formation oligosphaeridium patulum gleicheniidites conspiciendus nannoceratopsis gracilis mancodinium semitabulatum kekryphalospora distincta parvocysta spp. endoscrinium luridum gonyaulacysta jurassic glossodinium dimorphum, dichadogonyaulax pannea chasmatosporites hians quadraeculina anellaeformis mendicodinium reticulatum neoraistrickia gristhorpensis liasidium variabile dapcodinium priscum rhaetipollis germanicus rhaerogonyallax rhaetica illinites chitinoides enzonalasporites vigens vallasporites ignaciii camerosporites secatus, duplicisporites granulatus triadispora verrucata fig. 2. stratigraphical scheme showing upper tri assic and jurassic litho stratigraphic units, selected sequence strati graphic boundaries and selected last occurrence datums of selected taxa in the norwe gian–danish basin. rosa_2008:rosa-2008 01/07/09 15:48 side 22 the results of the correlation show that lithostratigraphic units and key sequence stratigraphic surfaces can be followed across the danish and norwegian areas indicating that the general depositional development along the transect largely follows the pattern described by nielsen (2003). hence, the sequence stratigraphic and lithostratigraphic schemes established for the danish part were successfully applied to the norwegian wells, as exemplified in figs 1, 2. however, the lack of readily recognisable sequence stratigraphic key surfaces and biostratigraphic events within the middle – upper triassic succession is a major problem for reliable stratigraphic analyses. the middle triassic – an unsuitable climate for palynomorph preservation during the middle triassic the norwegian–danish basin was situated around 35°n. the arid to semi-arid conditions that had prevailed during the early triassic continued, as signified by mainly fluvial and lacustrine, heterogeneous, siliciclastic rocks of the skagerrak formation (michelsen & clausen 2002). marine calcareous mudstones and carbonates of the muschelkalk formation in the north german basin ex panded northwards during a transgressive event in the anisian to early ladinian (michelsen & clausen 2002). arid to semi-arid climatic conditions continued during early late triassic times, with deposition of variegated, calcareous, anhydritic and pyritic mudstones in sabkhas and ephemeral lakes. in the deep central part of the basin more permanent lakes were established. the deposits are included in the oddesund (danish area) and smith bank (norwegian area) formations. towards the basin margins in the north and north-east these deposits pass into alluvial fans and fluvial sediments of the skagerrak formation. the combination of large lateral variations in depositional environment and the absence of extensive marine-flooding surfaces and marked unconformities hinder identification of reliable and regional sequence stratigraphic key surfaces. whereas middle to lower upper triassic successions of the arctic, alpine and tethys regions show evidence of fairly rich and diverse vegetation, the pre-rhaetian triassic redbeds of the norwegian–danish basin only contain sparse palyno morphs. however, this does not necessarily mean that there were no plants growing in the area at the time of deposition. most arid areas today host some vegetation adapted to such a hostile climate, but conditions for preservation of palynomorphs are generally poor due to oxidation of the sediments. nevertheless, the present study indicates that during a restricted interval in the latest middle triassic the area hosted a relatively rich vegetation and the climatic conditions in the area were better suited for palynomorph preservation. ladinian spore–pollen floras well-preserved, typical middle triassic palynofloras were found in ditch cuttings from the lowermost part of the oddesund formation in the felicia-1 well, and in the lower part of the skagerrak formation in k-1 and 11/10-1. these assemblages are distinguished from caved rhaetian–creta ceous material by their generally darker colour. all the assemblages contain illinites chitonoides, a pollen species that has a last appearance datum at the top of the ladinian in northern europe, and within the carnian in the arctic region (de graziansky et al. 1998). the assemblage from felicia-1 is dominated by the bisaccate pollen ovalipollis ovalis /pseudoalatus, which has its first common appearance datum at the base of the late ladinian (de graziansky et al. 1998). in addition, members of protodiploxypinus, e.g. p. fastidioides and p. macroverrucosus, and triadispora, mainly t. crassa, t. plicata, t. verrucata, are common constituents of the palynofloras. the co-occurrence of angustisulcites klausii, kuglerina meieri, podosporites amicus, staurosaccites quadrifidus, rimaesporites aquilonalis, ara tri sporites spp., camerosporites secatus, c. verrucatus, dupli ci sporites granulatus and enzonalasporites vigens also suggests a late ladinian age (de graziansky et al. 1998; schulz & heu nisch 2005). the presence in felicia-1 of the typical triassic chlorococcalean coenobium plaesiodictyon mosellanum with a known stratigraphical range from late anisian to latest norian indicates brackish to freshwater conditions. in the easternmost parts of the norwegian–danish basin comparable palynofloras are present in core samples from the höllviken-2 well and two other wells in southern sweden. the höllviken-2 palynoflora is dominated by monolete (ara trisporites spp.) and trilete (e.g. calamospora spp. and ana pi culatisporites spp.) spores, but bisaccate pollen are also abun dant and diverse. both abundant aratrisporites spores and species of protodiploxypinus suggest a latest muschelkalk – early keuper (i.e. ladinian) age, and this range is further limited to early keuper by the presence of retisulcites perforatus (last appearance datum in earliest carnian), nevesisporites lubricus, and ovalipollis brutus. a banana-shaped acritarch, dactylofusa sp., is present in the palynoflora from fcc-1. comparable palynofloral assemblages were found in well a-2 in the danish central graben containing e.g. aratri sporites saturni, angustisulcites klausii, illinites chitonoides, tria dispora spp., protodiploxypinus fastidioides, p. granulatus, pro to diploxy pinus spp., and striatoabieites aytugii (bertelsen 1975). thus, typical ladinian palynofloras appear to be present with in a relatively restricted time interval along the northern margin of the mid-triassic muschelkalk sea. 23 rosa_2008:rosa-2008 01/07/09 15:48 side 23 24 remarks on stratigraphy, environment and climate the höllviken-2 assemblages are recorded in a succession of dark, fine-grained sediments (maglarp-c member) with fossiliferous intervals containing fossil fish, ostracods and characean algae. previous biostratigraphy based on the latter correlates the succession with late muschelkalk to early keuper, i.e. mainly of ladinian age (kozur 1974). the characean algae suggest lim nic to possibly brackish environments, while the fish fauna vary from limnish-brackish (e.g. paleobates spp.) to fully marine (e.g. birgeria spp. and hybodus spp.). the ostracods (bairdia spp.) indicate warm, shallow littoral environments. hence, the combined macrofossil evidence suggests a shallow marine environment with brackish to limnic lagoons behind the coast. the palynoflora recorded in the höllviken-2 core supports a ladinian age. all the contemporaneous palynofloras described here contain several elements typical of warm and dry conditions. for instance, aratrisporites is a common constituent in arid to semi-arid triassic palynofloras worldwide. the parent plants, pleuromeiacean lycopsids, appear to have been opportunistic and saline-tolerant inhabitants of intertidal environments (retallack 1975). the taeniate bisaccates, e.g. lunatisporites and striatoabieites, are generally regarded to have been produced by pteridosperms, and adapted to warm and dry conditions. the bisaccate pollen of triadispora, protodiploxypinus, ovalipollis, illinites and staurosaccites, as well as the monosaccate pollen enzonalasporites, and pollen of the circumpolles group, i.e. duplicisporites and camerosporites, are all believed to come from conifers. the circumpolles group is only recorded in some of the investigated assemblages. they are regarded as relatives to the cheirolepids, a group of conifers often associated with warm and dry conditions. in palynofloras of similar age from central and south europe, pollen of the circumpolles group tend to be much more abundant. the diverse ladinian spore–pollen flora described above indicates a generally warm and dry climate, but the diversity and preservation suggest a change towards more humid conditions favouring preservation of palynomorphs during this interval. in germany and poland the uppermost ladinian, known as the lettenkeuper, is interpreted as reflecting more humid conditions. it seems plausible that the palynofloras recorded in this study reflect the same change to more humid conditions. conclusions the recorded ladinian palynofloras from the maglarp-c member in southern sweden, the lowermost oddesund formation in denmark, and from the middle skagerrak formation in both denmark and norway, enable correlation between these units in the otherwise poorly dated triassic succession in the norwegian–danish basin. the fact that the palynofloral assemblages are recognisable even in ditch cutting material of exploration wells makes them very useful markers. the palynoflora contains many elements indicating a warm and dry climate, and deposition in fresh to brackish water in coastal environments. the preservation of the palynoflora probably reflects a climatic event with a change to more humid conditions, similar to that recorded in northern and central europe. acknowledgement we thank talisman energy norway a/s for allowing us to publish results from the norwegian wells. references bertelsen, f. 1975: triassic palynology and stratigraphy of some danish north sea boreholes. danmarks geologiske undersøgelse årbog 1974, 17–32. de graziansky, p.-c., hardenbol, j., jacquin, t. & vail, p.r. (eds) 1998: mesozoic and cenozoic sequence stratigraphy of european basins. society for sedimentary geology (sepm) special publication 60, 786 pp. kozur, h. 1974: biostratigraphie der germanischen mitteltrias. freiberger for schungshefte c280 paläontologie, teil i, ii & anlagen, 56 pp. + 71 pp. michelsen, o. & clausen, o.r. 2002: detailed stratigraphic subdivision and regional correlation of the southern danish triassic succession. marine and petroleum geology 19, 563–587. nielsen, l.h. 2003: late triassic – jurassic development of the danish basin and the fennoscandian border zone, southern scandinavia. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 459–526. petersen, h.i., nielsen, l.h., bojesen-koefoed, j.a., mathiesen, a. & dalhoff, f. 2008: evaluation of the quality, thermal maturity and distribution of potential source rocks in the danish part of the nor wegian–danish basin. geological survey of denmark and greenland bulletin 16, 66 pp. retallack, g.j. 1975: the life and times of a triassic lycopod. alcheringa 1, 3–29. schulz, e., & heunisch, c. 2005: palynostratigraphische glieder ungs möglichkeiten des deutschen keupers. courier forschungsinstitut sen c kenberg 253, 43–49. authors’ addresses s.l., j.h.v., l.h.n., s.p., & k.d., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: sli@geus.dk m.e., geological survey of sweden, kiliansgatan 10, s-223 50 lund, sweden. rosa_2008:rosa-2008 01/07/09 15:48 side 24 geological survey of denmark and greenland bulletin 17, 2009, 65-68 the search for diamonds in greenland has resulted in the discovery of many new dykes of kimberlite and ultramafic lamprophyre and, most importantly, in the acquisition of a wealth of chemical data on rocks and minerals representing mantle material entrained by the dyke magmas. the discovery of a diamondiferous sheet at garnet lake in southern west greenland stimulated the research (hutchison 2005). over the past five to ten years, the geological survey of denmark and greenland together with the bureau of minerals and petroleum in greenland and international research groups have acquired, processed and interpreted data with the objective of identifying diamond-favourable regimes within the lithospheric mantle below the archaean craton in west greenland. here we present mineral data from drift samples that allow us to identify where mantle conditions in terms of lithology and depth may be favourable for the occurrence of diamonds. neoproterozoic igneous province the province comprises the sarfartoq carbonatite complex and abundant dykes and sills of carbonate-rich ultramafic silicate rocks (ultramafic lamprophyre and kimberlite) that have been emplaced in late neoproterozoic time into archaean rocks of southern west greenland between 65°n and 67°30´n (larsen & rex 1992; nielsen et al. in press; steenfelt et al. in press). magma emplacement was concentrated in the sarfartoq and maniitsoq regions (fig. 1), and took place from c. 600 ma to c. 555 ma (secher et al. in press). the first period of magmatism appears to be confined to the sarfartoq region while later magma pulses affected the entire region (fig. 1). a 568 ± 11 ma age of the garnet lake sheet (hutchison & heaman 2008) places this intrusion in the younger part of the period. ultramafic dykes (some of which are diamondiferous) of the same age in labrador, canada (tappe et al. 2006) show that the magmatism ex tended into the western part of laurentia, the then contiguous continent comprising north america and greenland. the magmatism is thought to have been triggered by incipient continental rifting. at about the same time, more prodiamonds and lithospheric mantle properties in the neo proterozoic igneous province of southern west greenland agnete steenfelt, sven monrad jensen, troels f.d. nielsen, karina k. sand and karsten secher © geus, 2009. geological survey of denmark and greenland bulletin 17, 65–68. available at: www.geus.dk/publications/bull 65 fig. 1. neoproterozoic igneous province and localities for rocks tested positive for dia monds (jensen et al. 2004, with later updates from intex resources asa, www.intexresources.com). the large diamond symbol marks the garnet lake commercial diamond operation. the sarfartoq carbonatite complex and dykes of kimberlite and ultramafic lamprophyre (uml) are emplaced into the archaean craton, which suffered deformation in the northern part during a palaeo proterozoic collision. southern boundary of palaeoproterozoic deformation maniitsoq region sarfartoq region 50°w 66°n 67°n 65°n 52°w54°w inland ice davis strait archaean craton qaqarssuk, 165 ma age 554–572 ma age 577–604 ma uml or kimberlite carbonatite diamond locality 50 km garnet lake sarfartoq greenland rosa_2008:rosa-2008 01/07/09 15:48 side 65 66 nounced rifting took place at the northern margin of laurentia and resulted in the intrusion of a prominent basaltic dyke swarm in north-west greenland and northern baffin land (dawes 2006). exploration and diamond discoveries exploration companies have used the so-called kimberlite indicator minerals in their search for host rocks for diamonds. samples of overburden or drift (mainly till) have been collected systematically over the entire archaean craton of southern west greenland and processed to obtain the nonmagnetic heavy mineral fraction, from which grains of mantlederived minerals including peridotitic garnet (pyrope), eclogitic garnet, chromite (chrome-spinel), picroilmenite and chrome-diopside have been picked under a microscope. many grains were subsequently chemically analysed to verify the visual identification and allow chemical classification. numerous samples with diamond-indicative, high-pressure mineral compositions indicate that the neoproterozoic province has a high prospective potential, and subsequent diamond tests have confirmed that the carbonate-rich ultramafic magmas brought up diamondiferous mantle at several localities within the province (fig. 1). the huge amount of mineral analyses were compiled and quality controlled by jensen et al. (2004). in 2004 hudson resources inc., guided by drift samples with an abundance of garnets derived from the diamond-stable mantle, discovered a significant amount of diamonds hosted in carbonatite-rich ultramafic rocks at garnet lake (fig. 1) in the sarfartoq region (hutchison 2005). con tinued exploration has established the presence of a 4 m wide, shallow-dipping sheet of kimberlitic rock with a promising diamond grade and diamond crystals up to 4 carats (0.8 g; hutchison & heaman 2008; www. hudsonresources.ca). harzburgite g10 400 600 800 1000 estimated garnet equilibrium temperature °c la tit ud e °n la tit ud e °n la tit ud e °n 1200 1400 1600 a b c graphite stable diamond stable 65 65.5 66 66.5 65 65.5 66 66.5 65 65.5 66 66.5 67 lherzolite g9 high-ti lherzolite g11 fig. 2. estimated temperatures of three classes (a: g9, b: g11 and c: g10) of mantle-derived garnet grains picked from the nonmagnetic, heavy mineral fraction of drift samples (mainly till). northern latitudes on vertical scale: upper group sarfartoq region (north of 66°n), lower group maniitsoq region. orange, open triangles are grains from the garnet lake area. red, vertical line marks the temperature of graphite-diamond phase transition (900°c). blue line (1200°c, c. 180 km depth) marks the cut-off value for deeply derived grains plotted in fig. 3. squares represent lithologically sorted temperature estimates based on mantle xenoliths (larsen & rønsbo 1993; garrit 2000; bizzarro & stevenson 2003; jensen et al. 2004; sand et al. in press); lines between ×-symbols are ranges in ni-in-garnet temperatures of grains in garnet concentrates (data from garrit 2000). large, filled, red triangles are temperatures determined on xenoliths from garnet lake (hutchison & heaman 2008). rosa_2008:rosa-2008 01/07/09 15:48 side 66 the neoproterozoic lithospheric mantle many mineralogical and chemical investigations of mantle xenoliths hosted by the kimberlites and ultramafic lamprophyres have demonstrated that the lithospheric mantle comprises an upper section of peridotitic rock types (lherzolite, harzburgite, dunite) depleted in elements such as ca, fe and ti relative to asthenospheric mantle because of extraction of large portions of basalt. the section of depleted mantle is underlain by a section with a predominance of fe-ti-rich, socalled fertile garnet lherzolite (references in fig. 2). it has also been demonstrated that some xenoliths from both regions have been derived from depths clearly within the high-pressure regime where diamond is stable (references in fig. 2). the constraints for the neoproterozoic geotherm have recently been improved to 38–41 mw/m2, and the thickness of the neoproterozoic lithosphere has been estimated to be at least 215 km over the entire province (sand et al. in press). studies by hutchison & heaman (2008) indicate that the diamonds at garnet lake probably formed at great depths within the fertile lherzolite, i.e. at temperatures above 1200°c, and within a period of 50 million years before the transporting mag ma brought them to the surface. the deep lithospheric mantle provenance of the xenoliths is also stressed by grütter & tuer (in press), who found an unusually high proportion of high-t perido titic garnets in drift samples from the immediate surroundings of garnet lake. garnets from deep lithospheric mantle garnet is the mineral that has been used most extensively in lithosphere studies and diamond exploration to reflect the temperature and pressure conditions as well as the lithology at the site where it equilibrated. the 15 000 available ana lyses of garnet grains from the neoproterozoic province therefore provide excellent material with which to locate dykes that have incorporated material with deep mantle provenance similar to that recorded at garnet lake. using a chemical discrimination system devised by grütter et al. (2004) we have selected garnets derived from depleted lherzolite (g9; fig. 2a), fertile lherzolite (g11, fig. 2b) and harzburgite (g10, fig. 2c), and determined their equi libration temperatures using mno concentrations (grüt ter et al. 1999). the results shown in fig. 2 are plotted against the latitude of the sample sites in order to reveal any regional differences. temperatures of garnet lake grains and published temperature estimates based on other minerals are shown for comparison. the diagrams show that a majority of the garnet grains derive from depths where diamond is the stable carbon phase, i.e. where the temperature is above 900°c. it appears that g11 garnets (from fertile lherzolite) mainly come from greater depths and have large populations over the entire latitude interval. this enforces the validity of current models invoking the ubiquitous presence of fertile lherzolite in the 67 fig. 3. neoproterozoic igneous province with localities of drift samples and results of screening garnet grain analyses belonging to classes g9, g10 and g11. deep garnets (green dots) have t-mn above 1200°c. the red symbols marking samples (drift or rock) with more than 10 deep garnets in the picked populations of peridotitic garnet grains show spatial correlation with diamond occurrences. southern boundary of palaeoproterozoic deformation maniitsoq region sarfartoq region 50°w 66°n 67°n 65°n 52°w54°w inland ice davis strait drift sample all sites with garnet with deep garnet with many deep garnets rock sample with many deep garnets carbonatite diamond locality 50 km garnet lake sarfartoq k rosa_2008:rosa-2008 01/07/09 15:48 side 67 68 lower lithospheric mantle section. a tendency for relatively more g11 grains above 1200°c in the maniitsoq region is observed. lherzolitic (g9) and harzburgitic garnets (g10) display origins in wider depth intervals, and very deep grains occur in both regions. the garnet lake garnets do stand out in reaching higher temperatures than many grains in the sarfartoq region. however, grains from several localities in the maniitsoq region have also yielded temperatures above 1400°c. it should be mentioned that the temperature is inversely correlated with mno concentrations, so that temperature data above 1600°c are uncertain owing to low analytical precision at low concentrations. the range in garnet t–mn temperature estimates is in good agreement with estimates using other methods, and the advantage of having the many additional data to establish a more statistically reliable, regional picture of mantle provenance is obvious. in addition, the drift-derived garnets provide information from areas where dykes have not been located or sampled. distribution of sites with high diamond potential figure 3 shows the extensive coverage of drift sample sites and the clusters of garnet-bearing samples (any mantlederived kind) where dykes are common (compare fig. 1). the deep (high-t) garnets have a narrower distribution, yet they are abundant in both regions. in order to highlight localities with a high proportion of deep garnets, an arbitrary lower limit of ten grains has been applied. rock-sample localities with a high proportion (more than ten grains) of deep garnets have been identified and are added as supplementary information. they outline additional localities with diamond potential in the sarfartoq region. the distribution of localities rich in deep garnet exhibits spatial correlation with that of diamond-bearing rocks and thus supports the observation made at garnet lake that an abundance of deep, lower lithospheric mantle material is characteristic of diamond-bearing dykes. however, the data also demonstrate that garnet lake is not unique in the province in this respect and the potential for making equally promising diamond discoveries elsewhere appears to remain. one small area near kangerlussuaq (k in fig. 3) has not yet proved positive for diamonds, but is considered a target for further exploration. subsurface exploration methods would be needed in that area, though, because poor exposure impedes surface recognition of significant dykes. references bizzarro, m. & stevenson, r.k. 2003: major element composition of the lithospheric mantle under the north atlantic craton: evidence from peridotite xenoliths of the sarfartoq area, southwestern greenland. contributions to mineralogy and petrology 146, 223–240. dawes, p.r. 2006: explanatory notes to the geological map of greenland, 1:500 000, thule, sheet 5. geological survey of denmark and green land map series 2, 97 pp. + map. garrit, d. 2000: the nature of the archaean and proterozoic lithospheric mantle and lower crust in west greenland illustrated by the geochemistry and petrography of xenoliths from kimberlites, 289 pp. unpublished ph.d. thesis, university of copenhagen, denmark. grütter, h. & tuer, j. in press: constraints on deep mantle tenor of sarfartoq-area kimberlites (greenland), based on modern thermobarometry of mantle-derived xenocrysts. lithos. grütter, h.s., apter, d.b. & kong, j. 1999: crust-mantle coupling: evidence from mantle-derived xenocrystic garnets. in: gurney, j.j. et al. (eds): proceedings of the viith international kimberlite conference 1, 307–313. cape town: red roof design. grütter, h.s., gurney, j.j., menzies, a.h. & winter, f. 2004: an updated classification scheme for mantle-derived garnet, for use by diamond explorers. lithos 77, 841–857. hutchison, m.t. 2005: diamondiferous kimberlites from the garnet lake area, west greenland: exploration methodologies and petrochemistry. danmarks og grønlands geologiske undersøgelse rapport 2005/68, 33–42. hutchison, m.t. & heaman, l.m. 2008: chemical and physical characteristics of diamond crystals from garnet lake, sarfartoq, west green land: an association with carbonatitic magmatism. the cana dian mineralogist 46, 1063–1078. jensen, s.m., secher, k., rasmussen, t.m. & schjøth, f. 2004: diamond exploration data from west greenland: 2004 update and revision. dan marks og grønlands geologiske undersøgelse rapport 2004/117, 90 pp. + 1 dvd. larsen, l.m. & rex, d.c. 1992: a review of the 2500 ma span of alkalineultramafic, potassic and carbonatitic magmatism in west greenland. lithos 28, 367–402. larsen, l.m. & rønsbo, j. 1993: conditions of origin of kimberlites in west greenland: new evidence from the sarfartoq and sukkertoppen regions. rapport grønlands geologiske undersøgelse 159, 115–120. nielsen, t.f.d., jensen, s.m., secher, k. & sand, k.k. in press: regional and temporal variations in the magmatism of the diamond province of southern west greenland. lithos. sand, k.k., waight, t., pearson, d.g., nielsen, t.f.d., makovicky, e. & hutchison, m.t. in press: the lithospheric mantle below southern west greenland: a geothermobarometric approach to diamond potential and mantle stratigraphy. lithos. secher, k., heaman, l.m., nielsen, t.f.d., jensen, s.m., schjøth, f. & creaser, r. in press: timing of kimberlite, carbonatite and ultramafic lamprophyre emplacement in the alkaline province located 64°– 67°n in southern west greenland. lithos. steenfelt, a., jensen, s.m., nielsen, t.f.d. & sand, k.k. in press: provinces of ultramafic lamprophyre dykes, kimberlite dykes and carbonatite in west greenland characterised by minerals and chemical components in surface media. lithos. tappe, s., foley, s.f., jenner, g.a., heaman, l.m., kjarsgaard, b.a., romer, r.l., stracke, a., joyce, n. & hoefs, j. 2006: genesis of ultramafic lamprophyres and carbonatites at aillik bay, labrador: a consequence of incipient lithospheric thinning beneath the north atlantic craton. journal of petrology 47, 1261–1315. authors’ addresses a.s, t.f.d.n. & k.s., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ast@geus.dk s.m.j., intex resources asa, munkedamsveien 45a, n-0250 oslo, norway. k.k.s., nano-science center, university of copenhagen, universitetsparken 5, dk-2100 copenhagen ø, denmark. rosa_2008:rosa-2008 01/07/09 15:48 side 68 geological survey of denmark and greenland bulletin 13, 2007, 49-52 49 the nordre strømfjord shear zone is a 1.8 ga zone of largescale, transcurrent and sinistral ductile shear (sørensen et al. 2006) within the nagssugtoqidian mobile belt (nmb) of central west greenland. it has been hypothesised that the nmb is a suture between two archaean continental masses (kalsbeek et al. 1987). during field work in 2005 along the nordre strømfjord shear zone in the fjord ataneq (fig. 1), some unusual rock types were discovered that preserve evidence of magmatic and metamorphic processes not previously reported in the area. these observations include the first indication of high-pressure (hp) metamorphism in west greenland and the first reported occurrence of a cumulate of giant orthopyroxene. the tectonic telescoping of these features together within the nordre strømfjord shear zone has important implications for reconstructing the palaeoproterozoic history of this region, and provides evidence that processes typical of phanerozoic continent–continent collision zones (e.g. the caledonian and alpine systems) operated at least as far back as 1.8 ga ago. high-pressure enclave on the north side of inner ataneq fjord an approximately 1.2 m wide and 4 m long lens of ultramafic rock occurs within strongly foliated garnet-sillimanite gneisses and schists, and garnet-bearing calc-silicate rock (fig. 2). the pale yellowish green, ultramafic rock is moderately foliated with its long axis parallel to the fabric in the enclosing gneisses. this core of the © geus, 2007. geological survey of denmark and greenland bulletin 13, 49–52. available at: www.geus.dk/publications/bull two tectonically significant enclaves in the nordre strømfjord shear zone at ataneq, central west greenland william e. glassley, john a. korstgård and kai sørensen fig. 1. north-eastern part of the agto map sheet (olesen 1984), with the localities of figs 2, 3 and 4 marked. the nordre strømfjord shear zone of the map area is characterised by vertically oriented supracrustal units alternating with quartzofeldspathic units, as also described by sørensen et al. (2006) from the area to the east. the regional amphibolite to granulite facies transition occurs over the eastern part of the map. 50 enclave consists of anthophyllite with a few minor additional phases. it is surrounded by a dark rim of dense, fine-grained rock approximately 30 cm thick (fig. 2) that is conformable to the shape of the ultramafic lens. the rim appears to be the result of a reaction between the silica-poor ultramafic rocks and the enclosing aluminiumand silica-rich metasediments, and it consists of olivine-orthopyroxene-clinopyroxene-spinelgarnet-amphibole. the fine-grained nature of the rim rock and its complex textural characteristics make it difficult to unambiguously decipher all aspects of its petrogenetic history. however, certain key observations show that the rim rock records an unusual history involving high-pressure metamorphism. the olivine occurs as remnant crystals that are occasionally seen to be in optical continuity but separated by pyroxene and spinel. garnets occur as isotropic areas that are nearly completely overgrown by spinel and pyroxene. garnet also occurs as inclusions in spinel. all combinations of grain-to-grain contacts have been observed, with the exception of garnet-olivine. there are also textural features suggesting that two generations of orthopyroxene and clinopyroxene may be present. these mineralogical features document a petrogenetic history in which the oldest mineral assemblage preserved in the rim of the enclave is garnet-olivine-orthopyroxene-clinopyroxene (i.e. garnet peridotite). the occurrence of garnet + olivine in ultramafic rocks and the occurrence of eclogite minerals in mafic compositions are the diagnostic mineral assemblages for hp metamorphism. defined in this way hp metamorphism is intermediate between granulite facies metamorphism and ultra high-pressure metamorphism (uhp) in which diamond and coesite are stable phases. in the hp enclave, the olivine + garnet-bearing assemblage is replaced, via reaction between olivine and garnet, by the assemblage spinel-orthopyroxene-clinopyroxene (i.e. spinel peridotite). olivine and garnet are preserved because the reaction was arrested before it went to completion. this metamorphism took place at a very low thermodynamic activity of water. replacement of a garnet peridotite mineral assemblage by that of spinel peridotite is the hallmark of recrystallisation during decompression from minimum pressures of about 18–20 kilobars (>60 km) and temperatures >750°c (schmä dicke & evans 1997; fumagalli & poli 2005). preliminary electron microprobe analyses of all of the mineral phases have been conducted. clinopyroxene-orthopyroxene geothermometry and orthopyroxene-garnet geobar ometry (brey & köhler 1990) intersect at 785°c and 21 kb. however, uncertainty in identifying cogenetic minerals, as well as the fact that these rocks have experienced extensive recrystallisation during decompression and cooling make it likely that these p–t conditions are a minimum; modifications are to be expected as further analyses are conducted. the electron microprobe data provide support for the argument that the high density rim around the ultramafic rock is, in fact, a metasomatic feature reflecting steep chemical potential gradients between the metasediments and the ultramafic rock. in particular, the very high modal abundance of the spinel (>20%) and the absence of detectable cr in any of the minerals are inconsistent with primary crystallisation from an ultramafic composition. rather, these characteristics suggest limited metasomatic reaction between the enclave and the surrounding metasediments into which we envisage the enclave to have been tectonically emplaced. giant orthopyroxene cumulate with inter stitial anorthosite and associated rocks approximately 3 km west of the hp site a series of gabbroic anorthosite and coarse-grained orthopyroxenite lenses occur that are metres to tens of metres in size (fig. 3). this series of lenses is traceable along the coast over a distance of 1 km. the fig. 2. lens (boudin?) of yellowish green, ultramafic rock within garnetbiotite-sillimanite gneiss and calc-silicate rock. the hammer (1 m) rests on the ultramafic rock and is just to the right of a dark, 30 cm thick rim (indicated by arrow) that completely encloses the ultramafic lens. the dark rim is the source of the garnet-spinel-olivine-orthopyroxeneclinopyroxene sample. for location see fig. 1. 51 margins of these lenses are tectonised at their contact with the enclosing quartzofeldspathic gneisses. the orthopyroxenites were observed in two distinct forms. one of these is a monomineralic lens of thumb-sized, equant, euhedral to subhedral orthopyroxene crystals. the lens is approximately two metres by four metres in size and exhibits no internal fabric. the other form is a spectacular giant orthopyroxene cumulate containing crystals more than 30 cm long and 15 cm wide that have a strong preferred orientation, with long axes parallel to each other in a classic cumulate texture. the crystals exhibit striking macroscopic kink banding (fig. 4). anorthosite is found as discontinuous films along the edges of the orthopyroxene crystals and as cuspate pockets where triple junctions of orthopyroxene crystals occur. in thin section the orthopyroxenites are seen to have preserved detailed evidence of a complex magmatic history and metamorphic recrystallisation, even though field evidence unequivocally shows these rocks to have been tectonically emplaced into their present setting. the primary magmatic mineral assemblage consists of remnant forsteritic olivine incompletely resorbed by orthopyroxene, green spinel and plagioclase with chromite, rutile, phlogopite, apatite and zircon as additional phases, either primary or a result of exsolution. all of these minerals are observed as inclusions within the orthopyroxene, as well as phases interstitial to orthopyroxene in pockets of anorthosite. secondary minerals associated with metamorphic recrystallisation are amphibole (as trains of small grains within orthopyroxene, occurring along crystallographically controlled planes) and quartz. reconnaissance electron microprobe analyses of the plagioclase show that its composition is affected by its environment: plagioclase grains within the anorthositic pockets are close to an60, while those contained within the orthopyroxene, which generally are associated with amphibole, are approximately an40. the amphibole is nearly pure cummingtonite. other observations made with the electron microprobe showed the presence of fe-ni sulphides and pure cu spherules. in addition, the spinels and phlogopites are tiand cr-rich. these characteristics of the orthopyroxenites suggest that the cumulates formed by gravitational settling of giant orthopyroxenes in a magma chamber. the presence of plagioclase, clinopyroxene and rutile exsolution lamellae suggests that the orthopyroxenes crystallised at high pressure (>10 kb), which is consistent with the co-existence of orthopyroxene-olivine-plagioclase-spinel. it has been postulated that anorthositic massifs form via fractionation of orthopyroxene from magmas of appropriate compositions at or near the base of the continental crust (emslie 1985). however, such cumulates have never been observed before, and the slivers of cumulate orthopyroxenite observed in ataneq may be the remnants of such a system that has been tectonically dismembered fig. 3. two 3 m long lenses of gabbro anorthosite approximately 2 km west of the ultramafic lens shown in fig. 2. note the duplex structure within the lens. for location see fig. 1. fig. 4. giant orthopyroxene crystals (dark olive green except where reflecting) separated by grey and white intercumulus anorthosite. glove above the large single crystal in the centre of the photograph is approximately 20 cm long. note kink banding in the central crystal (arrow points towards the kink band). opx, orthopyroxene. for location see fig. 1. conclusions the high-pressure rocks and orthopyroxenite cumulates observed in ataneq attest to tectonic telescoping of rocks that originated from profoundly different geological environments. they provide evidence that within the nordre strømfjord shear zone, samples of the deepest levels of continental crust and the upper mantle are present. these rocks occur within contrasting lithologies of the nordre strømfjord shear zone: the hp lens within a supracrustal unit (sensu sørensen et al. 2006) and the pyroxenite-anorthosite assemblage within a quartzofeldspathic gneiss unit. further to the east, a complex of lenses of ultramafic rocks, pillow lavas and unusual tourmaline-phlogopite rocks (interpreted to be the metamorphosed remnant of submarine hot-spring exhalations; sørensen et al. 2006) have been observed enclosed in supracrustal units. these rock types provide compelling evidence that upper mantle, deep continental crust and oceanic crust were tectonically juxtaposed, probably during continent–continent collision, later to be deformed within the nordre strømfjord shear zone. the tectonism responsible for the emplacement of these rocks within continental rocks may be thrust stacking as described by van gool et al. (1999) and sørensen et al. (2006) south of the shear zone near the inland ice. acknowledgement the work of w.e.g. was funded by the university of aarhus. references brey, g.p. & köhler, t. 1990: geothermobarometry in four-phase lherzolites. ii. new thermobarometers and practical assessment of existing thermobarometers. journal of petrology 31, 1353–1378. emslie, r.f. 1985: proterozoic anorthosite massifs. in: tobi, a.c. & touret, j.l.r. (eds): the deep proterozoic crust in the north atlantic provinces. nato asi series 158, 39–60. fumagalli, p. & poli, s. 2005: experimentally determined phase relations in hydrous peridotites to 6.5 gpa and their consequences on the dynamics of subduction zones. journal of petrology 46, 555–578. kalsbeek, f., pidgeon, r.t. & taylor, p.n. 1987: nagssugtoqidian mobile belt of west greenland: a cryptic 1850 ma suture between two archaean continents – chemical and isotopic evidence. earth and planetary science letters 85, 365–385. olesen, n.ø. 1984: geological map of greenland, 1:100 000, agto, 67 v.1 nord. copenhagen: geological survey of greenland. schmädicke, e. & evans, b.w. 1997: garnet-bearing ultramafic rocks from the erzgebirge, and their relation to other settings in the bohemian massif. contributions to mineralogy and petrology 127, 57–74. sørensen, k., korstgård, j.a., glassley, w.e. & stensgaard, b.m. 2006: the nordre strømfjord shear zone and the arfersiorfik quartz diorite in arfersiorfik, the nagssugtoqidian orogen, west greenland. geological survey of denmark and greenland bulletin 11, 145–161. van gool, j.a.m., kriegsman, l.m., marker, m. & nichols, g.t. 1999: thrust stacking in inner nordre strømfjord area, west greenland: significance for the tectonic evolution of the palaeoproterozoic nagssugtoqidian orogen. precambrian research 93, 71–86. authors’ addresses w.e.g. & j.a.k., geological institute, university of aarhus, høegh-guldbergsgade 2, dk-8000 århus c, denmark. e-mail: geobg@nf.au.dk k.s., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. 52 ultramafic dyke rocks with kimberlitic megacrysts and mantle nodules have been known for decades from the northern part of the archaean block and adjacent proterozoic terranes in southern west greenland (fig. 1; escher & watterson 1973; goff 1973; scott 1981; larsen & rex 1992; mitchell et al. 1999). some of the dykes have proved to be diamondiferous (see jensen et al. 2004a, b, for exploration results, diamond contents, and references). the c. 600 ma old dykes were called ‘kimberlitic’ by larsen & rex (1992), but mitchell et al. (1999) concluded that they were best referred to a ‘carbonatiteultramafic lamprophyre’ suite (aillikites or melnoites). mitchell et al. (1999) further suggested that the west greenland province represents “one of the few bona fide examples of ultramafic lamprophyre which contain diamonds”. reports on indicator mineral assemblages (jensen et al. 2004b) and diamond contents (e.g. hudson resources inc. 2005) have re-opened the discussion on the classification of the dykes. the results of an investigation of the majuagaa dyke (nielsen & jensen 2005) are summarised below, together with the preliminary results of a regional investigation of the groundmass minerals of the dykes. it is concluded that dykes in the maniitsoq region are similar to archetypal, south african, on-craton, type 1 kimberlites, and that all regions of the west greenland province of ultramafic magmatism are favourable for diamond exploration. the majuagaa dyke the majuagaa dyke (jensen et al. 2004a) is 2.5 km long and up to 2 m wide. it is located c. 50 km sse of maniitsoq (fig. 1) and strikes wsw–ene. the dyke is dark grey with many olivine-rich fragments (up to 10 cm) and rounded megacrysts of ilmenite (up to 4 cm). it contains the classic kimberlitic suites of megacrysts and mantle nodules, including eclogite (jensen & secher 2004, fig. 5). the groundmass is fine-grained and composed of olivine fragments, calcite, serpentine, ilmenite and minor mg-rich spinel. phlogopite and apatite are rare. the dyke is diamondiferous (jensen et al. 2004a). samples were collected along the length of the dyke. sixty thin sections (fig. 2) were examined and a number selected for an electron microprobe study. all mineral data from groundmass, megacysts and nodules, the bulk chemistry, and analytical techniques are reported in nielsen & jensen (2005). classification of the majuagaa dyke mitchell (1995) and tappe et al. (2005) use the following criteria for the classification of kimberlite (s.s.): (1) the groundmass contains no clinopyroxene; (2) groundmass spinel belongs to the magmatic trend 1 (mg-rich titanomagnetite); (3) phlogopite is zoned towards the al2o3and bao-rich kinoshitalite endmember and (4) ilmenite has a high geikilite component (> 40 mol.% mgtio3) and little pyrophanite (mntio3). mitchell et al. (1999) found that these criteria were not met by the west greenland dykes and concluded they were ultramafic lamprophyres (aillikites or melnoites) rather than kimberlites. archetypal kimberlite from the maniitsoq region, southern west greenland and analogy to south africa troels f.d. nielsen, martin jebens, sven m. jensen & karsten secher © geus, 2006. geological survey of denmark and greenland bulletin 10, 45–48. available at: www.geus.dk/publications/bull sisimiut sisimiut maniitsoq maniitsoq kangerlussuaq kangerlussuaq fjord qaqqaarsuk tupertalik majuagaa sarfartoq sarfartoq 50 km kimberlitic occurrence carbonatite complex diamond occurrence southern boundary of palaeoproterozoic deformation 65°n 66°n 67°n 50°w51°w53°w 52°w greenland fig. 1. kimberlites and ultramafic lamprophyres (kimberlitic occurrences), carbonatite complexes and diamond occurrences in southern west greenland (after jensen & secher 2004). regions are indicated. 45 nielsen & jensen (2005) made the following observations in the majuagaa dyke: • the clinopyroxene criteria: no clinopyroxene was found in the groundmass. • the magmatic trend 1 spinel criteria: the cores of euhedral spinel grains (< 0.1 mm across) have compositions in the magmatic trend 1 field (fig. 3). mg-rich rims compare with spinels of south african calcite-kimberlite (mitchell et al. 1999). • the ilmenite criteria: groundmass grains conform with the compositions from archetypal kimberlite (fig. 4), whereas megacrysts appear to be xenocrystic (nielsen & jensen 2005). • the phlogopite criteria: tiny, euhedral, clear to weakly greenish flakes are rich in al2o3 and bao (fig. 5), poor in tio2 and feo (total) and rich in the kinoshitalite end member. they conform with phlogopite of archetypal kimberlite (see mitchell 1995). majuagaa bulk composition the bulk composition of the majuagaa dyke is kimberlitic (see nielsen & jensen 2005). the average ree (fig. 6) and trace element (fig. 7) compositions of the majuagaa dyke folow the base of the fields of kimberley (south africa), oncraton, type 1 kimberlites (le roex et al. 2003). the majuagaa dyke shows positive ti-, nband ta-anomalies. they are caused by a high proportion of ilmenite megacrysts. in mineralogy (see above) as well as bulk chemistry the majuagaa dyke is best compared to classic south african, oncraton, type 1 kimberlite. 46 0.2 0.0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1.0 magmatic trend t1 0.3 0.4 0.5 0.6 fe t 2+/(fe t 2++mg) 0.7 0.8 0.9 1.0 t i/ (t i+ c r+ a l) magmatic trend t2 orangeite trend lamproite trend fig. 2. plain polarised photograph of thin section (22 x 40 mm) from the majuagaa kimberlite. the rock is composed of fragments of olivine, ilmenite and pyroxene from disintegrated megacrysts and nodules in a groundmass of calcite, serpentine, spinel and ilmenite and rare phlogopite and apatite. only a single olivine microphenocryst is observed (p). m: olivine megacryst; px: fragment of clinopyroxene megacryst and ilm: ilmenite megacryst. fig. 3. cores of groundmass spinels plot in the field of the magmatic trend 1 spinels of archetypal kimberlites (after nielsen & jensen 2005). rim compositions to the left of the field are characteristic of calcite kimberlites (mitchell 1995). all other fields after mitchell (1995). k p geikielite mgtio 3 pyrophanite mntio 3 ilmenite fetio 3 100 80 60 40 20 0 0 20 40 60 80 100 0 20 40 60 80 100 fig. 4. proportions of ilmenite endmembers of the majuagaa dyke. fields for kimberlite and premier mine kimberlite (p) from south africa are shown (after nielsen & jensen 2005). regional variations the west greenland province is part of the c. 600 ma old north atlantic province of carbonatite and ultramafic alkaline magmatism, from the torngat region (canada) to the archaean of west greenland (tappe et al. 2004). the compositions of groundmass phlogopite reflect the compositions of the melts (mitchell 1995). the preliminary results of a regional investigation (fig. 8) suggest a gradual evolution from oncraton, south african type 1 kimberlite in the west greenland archaean craton (maniitsoq), through a kimberlite/ultramafic lamprophyre (kimberlite/aillikite) zone at the border between archaean and proterozoic terranes (sarfartoq and along the kangerlussuaq fjord, fig. 1) to ultramafic lamprophyre (aillikite/melnoite) magmatism in the proterozoic terranes of sisimiut (greenland; scott 1977) and torngat (canada; tappe et al. 2004). the diamond potential results of hudson resources inc. (2005) suggest that stones of gem quality and size may be found in west greenland. nevertheless, it appears to be an issue for some exploration companies and investors that mitchell et al. (1999) classified the west greenland dykes as ultramafic lamprophyres and implied that true kimberlite was not found. however, the majuagaa dyke documents that diamondiferous, archetypal type 1 kimberlite occurs in the west greenland province. hutchison (2005) describes the best investigated and most promising west greenland diamond occurrence at ‘garnet lake’ (border zone; sarfartoq region, fig. 1). the ‘garnet lake’ dykes have characteristics of both kimberlite and ultramafic lamprophyre and have features reminiscent of south african orangeite. finally, the ultramafic lamprophyres (aillikites/melnoites) of the proterozoic sisimuit region compare with diamondiferous ultramafic lamprophyres (aillikites) of the torngat region (fig. 8; tappe et al. 2004), and a diamond potential is also indicated in the little prospected sisimiut region. conclusions the c. 600 ma old ultramafic magmatism of the west greenland province shows – from the archaean craton to the proterozoic terranes – a transition from classic south african, on-craton type 1 kimberlite to ultramafic lamprophyre (aillikite/melnoite). diamonds are recovered from the entire range and a diamond potential thus exists throughout the west greenland province. 47 0 0 1 2 3 4 5 2 4 6 8 rim core b ao ( w t % ) tio 2 (wt %) fig. 5: bao vs. tio2 in groundmass phlogopite of the majuagaa dyke showing the general increase in bao and decrease in tio2 in the margins of the grains (after nielsen & jensen 2005). group i kimberlites la ce pr nd smp eu gd tb dy ho er tm yb lu r o ck /c h o n d ri te 0 10 100 1000 on-craton group i kimberlites on-craton rb r o ck /p ri m it iv e m an tl e 0 1 10 100 1000 bath u k tanb lacepb pr sr nd p smhf zr eu ti gdtbdyhoertmyb lu fig. 6. chondrite normalised bulk ree concentrations of the majuagaa dyke (after nielsen & jensen 2005). grey field: kimberley (south africa), type 1 kimberlites (le roex et al. 2003). fig. 7. trace element concentrations of the majuagaa dyke normalised to primitive mantle (after nielsen & jensen 2005). grey field: kimberley (south africa), type 1 kimberlites (le roex et al. 2003). acknowledgements the investigation was carried out under contract with the bureau of minerals and petroleum (bmp), nuuk, greenland. references escher, a. & watterson, j. 1973: kimberlites and associated rocks in the holsteinsborg – søndre strømfjord region, central west greenland. rapport grønlands geologiske undersøgelse 55, 26–27. goff, s.p. 1973: the mineralogy and mineral chemistry of a kimberlite dike from søndre isortoq fjord, south-west greenland, 103 pp. unpublished ph.d. thesis, university of leicester, united kingdom. hudson resources inc. 2005: hudson finds larger diamonds at garnet lake and confirms new diamond area. web release, december 1, 2005 (info@hudsonresources.ca). hutchison, m.t. 2005: diamondiferous kimberlites from the garnet lake area, west greenland: exploration, methologies and petrochemistry. in: secher, k. & nielsen, m.n. (eds): workshop on greenland´s diamond potential, 7–9 november 2005, in copenhagen. danmarks og grønlands geologiske undersøgelse rapport 2005/68, 33–42. jensen, s.m. & secher, k. 2004: investigating the diamond potential of southern west greenland. geological survey of denmark and greenland bulletin 4, 69–72. jensen, s.m., secher, k. & rasmussen, t.m. 2004a: diamond content of three kimberlitic occurrences in southern west greenland. diamond identification results, field description and magnetic profiling. danmarks og grønlands geologiske undersøgelse rapport 2004/19, 41 pp. jensen, s.m., secher, k., rasmussen, t.m. & schjøth, f. 2004b: diamond exploration data from west greenland: 2004 update and revision. danmarks og grønlands geologiske undersøgelse rapport 2004/117, 90 pp. (1 dvd enclosed). larsen, l.m. & rex, d.c. 1992: a review of 2500 ma span of alkaline, ultramafic, potassic and carbonatitic magmatism in west greenland. lithos 28, 367–402. le roex, a.p., bell, d.r. & davis, p. 2003: petrogenesis of group 1 kimberlites from kimberley, south africa: evidence from bulk-rock geochemistry. journal of petrology 31, 779–812. mcdonough, w.f. & sun, s.-s. 1995: the composition of the earth. chemical geology 120, 223–253. mitchell, r.h. 1995: kimberlites, orangeites and relates rocks, 410 pp. new york / london: plenum press. mitchell, r.h., scott smith, b.h. & larsen, l.m. 1999: mineralogy of ultramafic dikes from the sarfartoq, sisimiut and maniitsoq areas, west greenland. in: gurney, j.j. et al. (eds): proceedings of the vita international kimberlite conference 2, 574–583. cape town: red roof design. nielsen, t.f.d. & jensen, s.m. 2005: the majuagaa calcite-kimberlite dyke, maniitsoq, southern west greenland. danmarks og grønlands geologiske undersøgelse rapport 2005/43, 59 pp. scott, b.h. 1977: petrogenesis of kimberlites and associated potassic lamprophyres from central west greenland, 133 pp., + 6 appendices. unpublished ph.d. thesis, university of edinburgh, united kingdom. scott, b.h. 1981: kimberlite and lamprophyre dykes from holsteinsborg, west greenland. meddelelser om grønland. geoscience 4, 24 pp. tappe, s., jenner, g.a., foley, s.f., heaman, l.m., besserer, d., kjarsgaard, b.a. & ryan, a.b. 2004: torngat ultramafic lamprophyres and their relation to the north atlantic alkaline province. lithos 76, 491–518. tappe, s., foley, s.f., jenner, g.a. & kjarsgaard, b.a. 2005: integrating ultramafic lamprophyres into the iugs classification of igneous rocks: rational and implications. journal of petrology 46, 1893–1900. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: tfn@geus.dk 48 phlogopite compositions, regional 0 1 2 3 4 5 6 7 8 tio2 wt% a l 2 o 3 w t% sarfartoq kangerlussuaq maniitsoq sisimiut main torngat torngat kimberlite 0 5 10 15 20 fig. 8. al2o3 vs. tio2 of groundmass phlogopites from the maniitsoq (archaean craton), sarfartoq and kangerlussuaq fjord (border zone to the proterozoic) and sisimiut (proterozoic) regions. the sisimiut data (scott 1977) includes one locality referred to the border zone. kimberlite box in top-left corner after mitchell (1995). torngat field after tappe et al. (2004). until recently, in situ u-pb zircon geochronology could be carried out only using ion microprobes, requiring lengthy analysis times of c. 20 minutes. however, new developments in laser ablation inductively coupled plasma mass spectrometer technologies have resulted in zircon geochronology techniques that are much faster, simpler, cheaper, and more precise than before (e.g. frei et al. 2006, this volume). analyses approaching the precision obtained via ion microprobe can now be undertaken in 2–4 minutes using instruments such as the 213 nm laser ablation (la) system coupled with element2 sector-field inductively coupled plasma mass spectrometer (sf-icp-ms) housed at the geological survey of denmark and greenland (geus). the up to tenfold decrease in analytical time means that zircon geochronology can now be used in a much wider range of studies. the godthåbsfjord region, southern west greenland, contains some of the oldest rocks exposed on the earth’s surface reflecting a very complex archaean geological evolution (figs 1, 2). over recent years geus has undertaken a range of mapping projects at various scales within the godthåbsfjord region (see also below). these include the mapping of the 1:100 000 scale kapisillit geological map sheet (fig. 1), and regional and local investigations of the environments of formation and geological evolution of supracrustal belts, hosting potentially economic mineral occurrences. zircon geochronology is an important tool for investigating a range of geological problems in this region. by breaking down the complex geology into a series of simple problems that can be addressed using this tool, the geological evolution can be unlocked in a stepwise manner. three examples are presented below: (1) the mapping of regional structures; (2) characterising and correlating supracrustal belts; and (3) dating metamorphism and mineralisation. although focus is on the application of zircon geochronology to these problems, it is important to note that the resulting data must always be viewed within a wider context incorporating geological mapping and structural, geochemical and petrographic investigations. regional geology the geology of the godthåbsfjord region is dominated by orthogneiss formed during several distinct episodes of crustal growth during the archaean (fig. 2). these different-aged gneisses are thought to represent distinct small continental blocks that were amalgamated during the neoarchaean (at © geus, 2006. geological survey of denmark and greenland bulletin 10, 49–52. available at: www.geus.dk/publications/bull using zircon geochronology to resolve the archaean geology of southern west greenland julie a. hollis, dirk frei, jeroen a.m. van gool,adam a. garde and mac persson 49 fig.3 c. 3.87-3.6 ga c. 3.0-2.98 ga c. 3.2-2.98 ga c. 3.87-3.6 ga & c. 2.835-2.75 ga c. 2.92-2.84 ga greenland supracrustal belts terrane boundary 25 km nuuk st or ø kapisillit 64°n 51°w52°w 50°w 64°30' n godthåbsfjord ameralla ameralik fig. 1. overview map of the godthåbsfjord region. inset shows the location of the main map in greenland. supracrustal belts are shown in green. the boundaries of the 1:100 000 scale kapisillit geological map sheet area are in red. blue lines outline the area shown in fig. 3. bold black lines are inferred terrane boundaries, with the major age components of the different terranes indicated (after friend & nutman 2005). geological survey of denmark and greenland bulletin 17, 2009, 73-76 meeting the technical challenges posed by the arctic environment is a key issue in the development of greenland’s economy, particularly in the light of increasing interest in developing greenland’s mineral resources both onand offshore. this paper describes some results of the glaciological investigations carried out at malmbjerg. the world-class malmbjerg molybdenum prospect (71°50´n, 24°16´w) is located in the stauning alper, at the confluence of arcturus gletscher and the larger schuchert gletscher (fig. 1). the mineral occurrence was discovered in the 1950s and industrial development of the site was attempted in the following decades (henriksen 2008), but the venture remained economically unviable, largely due to its geographical situation. however, recent technological advances and high demand on the international metal market have led to revived interest in the development of the site as an open-pit mine. the two glaciers surrounding malmbjerg are pivotal in the planning of the mine. in particular, surface access to the mine across arcturus gletscher is a prerequisite, and the effects of accumulating coarse-grained mine waste on the glacier along the eastern slope of malmbjerg need to be investigated. load ing of the glacier by the weight of waste rock will modify the ice flow, and darkening of the glacier surface by rock dust will affect meltwater production and, through differential ablation, the elevation and morphology of the glacier surface. clean glaciers that become covered by debris from naturally occurring rock falls exhibit a much thinner and more homogeneously distributed debris cover, but may nevertheless provide data for comparison. finally, surge-type glaciers are known to exist in stauning alper (fig. 1), and the likelihood of a surge of arcturus and schuchert gletschers during the anticipated lifetime of the mine needs to be assessed. the results discussed below are based on data from an automatic weather station (aws) set up on schuchert gletscher in april 2008 by the geological survey of denmark and greenland (geus), and on field observations and groundpenetrating radar surveys carried out in september 2008. meteorological observations and melt modelling an estimate of the magnitude and regime of surface meltwater production and of ablation over the glacier surface is requi red to properly dimension the mine’s infrastructure, to model future differential ablation under a thickening cover of dust and debris and to assess the likelihood of surges of the two glaciers. obser vations from the aws on the glacier surface allow modelling of the surface energy ba lance and quantifying ablation. meltwater from the winter snow cover is important since it contributes to the total surface runoff from arcturus glet scher. a snow pit showed 576 mm water equivalent of accumulation (corresponding to winter balance) in early april 2008 at the aws site, where the snow cover had completely melted by mid-june. 73 glaciological investigations at the malmbjerg mining prospect, central east greenland michele citterio, ruth mottram, signe hillerup larsen and andreas ahlstrøm © geus, 2009. geological survey of denmark and greenland bulletin 17, 73–76. available at: www.geus.dk/publications/bull fig. 1. topographic map of the malmbjerg area. contour interval 100 m. the red circle shows the location of the planned mine. inset map shows the location of malmbjerg in east greenland. stauning alper is located immediately to the south-west of malmbjerg. " 0. schuchert gletscher ar ct ur us g let sc he r 2 km malmbjerg 72°n 24°20’w greenland gpr survey line gpr line shown in fig. 3 contact between arcturus and schuchert gletschers profile used for glacier flux calculation ablation stake automatic weather station 2008 rosa_2008:rosa-2008 01/07/09 15:48 side 73 the aws records air temperature and humidity in an aspirated radiation shield, as well as wind speed and wind direction, incoming and outgoing shortwave and longwave radiation, barometric pressure, snow depth, ice ablation and the ice-temperature profile at eight levels down to 10 m below the glacier surface (fig. 2). the measured data are stored locally in the data-logger memory and are transmitted to geus in copenhagen via a satellite link. the system is mounted on an aluminium tripod standing freely on the ice surface, so that the sensor height above the ice surface remains constant throughout the ablation season. aws data from may to august 2008 and topographic grids were used as input to the surface energy balance model of hock & holmgren (2005) with hourly time-steps over a grid of 50 × 50 m cells. figure 3a shows a good match between modelled and measured ablation at the aws site on arcturus gletscher, where the surface roughness length for wind speed over ice was set to 1 cm in order to best simulate the on-site ablation measured from the snow pit. since no field measurements of air temperature lapse rate or accumulation gradient were available from the area, the model was run for a range of values in order to assess the sensitivity of the mo del to these parameters (values of –3.5, –0.50 and –0.65°c/km and 0.0, 0.5 and 1.0 kg/m2/m of elevation/year were used). figure 3b shows the modelled cumulative ablation from may to august 2008 over the entire mod el led area with a lapse rate of –5.0°c/km and an accumulation gradient of 0.5 kg/m2/m of elevation/year. abundant snowfall at the end of august effectively ended any further significant ablation in 2008. the amount of meltwater produced in 2008 can then be deduced by integrating the model output over the glacier area of interest. effects of future glacier surface darkening deposition of rock dust derived from blasting and handling of ore and rock waste will result in darkening of the glacier surface around malmbjerg, with local consequences for mo bil ity across arcturus gletscher. dust accumulation reduces albedo and initially results in significantly higher ablation rates. ablation reaches a maximum when the effective thickness of the debris layer is attained, and then gradually decreases until the same ablation rate as clean, bare ice is observed at the critical thickness. any further thickening of the debris layer results in lower ablation rates, which become negligible under several tens of centimetres or a few metres. østrem (1959) first established an empirical relationship between ablation rate and thickness of the debris cover (the ‘østrem curve’). observed values for the effective and critical thicknesses range within 0.25–10 mm and 1.33–30 mm, respectively (kirk bride & dugmore 2003), with debris lithology and local climate playing significant roles. for simplicity we assumed in our model that dust 74 10 9 8 7 11 25 6 14 3 12 fig. 2. the automatic weather station: 1: as pi rated radiation shield for air temperature and humidity probes. 2: radiometers for shortwave and longwave incoming and outgoing radia tion. 3: sensor for wind speed and direction. 4: sensor to measure snow surface level. 5: two-axes tilt sensor of the instruments boom. 6: iridium satellite antenna. 7: iridium satellite radio. 8: atmospheric pressure. 9: sensor to meas ure ice-surface level. 10: thermistor string for ice temperature (drilled into the ice). 11: ablation sensor. 12: gps antenna. rosa_2008:rosa-2008 01/07/09 15:48 side 74 deposition will decrease exponentially with distance from the source. ice ablation is scaled depending on local debris thickness using an assumed østrem curve, converted into ice-thickness change and subtracted from the digital elevation model of the preceding time step. ice dynamics are neglected on the assumption that the glacier is in steady state during the entire modelled time, i.e. emergence and submergence velocity fields remain constant. this is a significant simplification, especially since glacier dynamics will respond to the modified ablation rates on the darkened surface. however, errors due to non-steady state were assumed to vary slowly with distance on the glacier surface, and to be of tolerable magnitude given the semi-quantitative character of this analysis. modelled results from various dust dispersal patterns show that areas with thick debris cover will become progressively raised over surrounding cleaner ice, and that the total surface meltwater production will peak after a variable number of years; from then on it will fall towards present values. ice-radar survey and surge potential a glacier surge is a strong and comparatively short-lived acceleration of ice flow lasting from a few months to a few years. it is typical of some glaciers with ice dynamics that oscillate without converging to a steady state in which accumulation in the higher parts is balanced by ice flow and ablation in the lower tongue. in the lower parts of a glacier a surge may result in the ice velocity increasing by one or even two orders of magnitude, with the surface becoming heavily crevassed and the glacier terminus possibly advancing by several kilometres. during the much longer quiescent phase following a surge, the ice at lower elevations will stagnate and waste down, while new mass builds up in the accumulation zone until the next surge event. a surge of either arcturus gletscher or schuchert gletscher would cause considerable difficulties to the mining operations. surge-type glaciers are known from stauning alper (olesen & reeh 1969; jiskoot et al. 2003), and hummocky morphology with extensive ice-cored moraines suggest that both the arcturus gletscher and schuchert gletscher may be of surge type. to assess the surge potential of arcturus gletscher, we estimated how the present mean-ice velocity through a given cross-section of the glacier compares to the balance velocity required in a steady state. profiles of arcturus gletscher were obtained using groundpenetrating radar (fig. 4). processing included mi gration to produce a geometrically correct representation of the subsurface. two other features could also be mapped: the cold to temperate transition surface separating cold ice from the underlying temperate ice at the pressure melting point and an englacial meltwater channel. analyses based on the balance velocity have commonly compared the ice flow through a cross-section to the net balance up-flow of the cross-section, but no accumulation data are yet available for the two glaciers. we therefore compared the ice flux with the net balance down-flow of the selected cross-section. although conceptually analogous, this method is not optimal for our purposes, since it cannot detect any mass build-up in the higher reaches of the glacier which are not balanced by the present glacier flow. however, melvold & hagen (1998) observed the imbalance between mass supply by ice flow and surface net balance to be strong along the entire length of a surge-type glacier in its quiescent phase, and this adds confidence to our approach. the balance velocity calculated for arcturus gletscher through the section 75 fig. 3. a: measured versus modelled mass balance from april to august 2008 at the site of the automatic weather station. the last winter snow melted around day 178. b: modelled ablation of arcturus gletscher and the lower part of schuchert gletscher. contour interval 100 m. 150 160 170 180 190 200 210 220 230 julian days –300 –250 –200 –150 –100 –50 0 m as s ba la nc e (c m w at er e qu iv al en t) measured modelled high: 3.6 low: 0 a bbb modelled melt (m water equivalent) 20 m bedrock surface transition between cold and temperate ice 200 m englacial meltwater channel fig. 4. migrated ground-penetrating radar profile of arcturus gletscher along the path shown in fig. 1. rosa_2008:rosa-2008 01/07/09 15:49 side 75 marked in fig. 1 is 18 m per year. this can be compared with the present-day average velocity of 22 m per year obtained from feature tracking on two orthophotographs from 2005 and 2007, and using theoretical results relating surface velocity with mean velocity over a cross-section of parabolic shape (paterson 1994). considering the uncertainties involved, we conclude that no imbalance has been detected to suggest that a surge of arcturus gletscher will occur in the near future. however, instead of a single year of observations, multi-year average net mass-balance data are needed to provide a more solid assessment. accumulation of waste rock and ice deformation waste-rock accumulation on the surface of arcturus gletscher close to the malmbjerg slope will affect the dynamics of the glacier, but prediction of such changes and their role is a difficult task, which requires sophisticated three-dimensional numerical models. as a preliminary approach we used simple oneand two-dimensional models to calculate the rate of ice deformation under a specific load, and how this impact could be minimised. a possible similar case may be found at the kumtor mine in the tien shan mountains in kyrgyzstan. gold-mine waste deposited on the davidov glacier over a period of 15 years caused a narrowing of the effective flow area, leading to an acceleration of the confined flow unit and increased crevassing extending to the adjacent flow unit of the glacier (bruce et al. 2008). in the first of our simplified conceptual models we assume the glacier to be a very viscous liquid, and that the rock waste, being thicker than naturally occurring supraglacial debris, will slowly sink into the glacier at a rate controlled by the viscosity of this liquid and the dimensions and shape of the rock particles. this model is analogous to the classical high-school experiment of dropping a ball into syrup and measuring how quickly it settles in order to determine the viscosity of the fluid. the second model assumes that all ice is displaced by the rock, and that the rate of deformation depends on the applied stress, employing a non-linear flow law commonly used in numerical models of glaciers. the two simplified models have been run using the expected timing and pattern of waste-rock accumulation, the profile of the glacier bed rock, and the depth to the temperature transition surface. several assumptions are involved, for instance concerning the applied stresses and the ice viscosity. nevertheless, the two models can provide rough upper and lower limits to what is likely to happen and how quickly. both models predict the highest settling velocities at the lower end of the deposit where the ice is thickest, and the fastest times to the bed (from less than a decade to a few decades) close to the malmbjerg slope where the ice is thinner. final remarks planning mining operations at malmbjerg must include an evaluation of the challenges posed by the glaciers that surround it. glaciological investigations can provide useful insight into how such challenges can be dealt with as the site is being developed. at the same time, mining sites such as malmbjerg can become case studies of great scientific interest. however, a single year of observations may not be representative of the local climate; therefore continued operation of the aws and further surveys are required to provide a better basis for validation of the results. acknowledgements this research has been funded by quadra mining ltd. jesper kofoed from quadra mining and the team from mt højgaard a/s at malmbjerg base camp helped with all aspects of our field work. references bruce, i., redmond, d. & thalenhorst, h. 2008: technical report on the 2007 year-end mineral reserves and resources, kumtor gold mine, kyrgyz republic, for centerra gold inc and cameco corporation, 162 pp. toronto: strathcona mineral services limited. henriksen, n. 2008: geological history of greenland, 270 pp. copen hagen: geological survey of denmark and greenland. hock, r. & holmgren, b. 2005: a distributed surface energy-balance model for complex topography and its application to storglaciären, sweden. journal of glaciology 51(172), 25–36. jiskoot, h., murray, t. & luckman, a. 2003: surge potential and drainagebasin characteristics in east greenland. annals of glaciology 36, 142–148. kirkbride, m.p. & dugmore, a.j. 2003: glaciological response to distal tephra fallout from the 1947 eruption of hekla, south iceland. journal of glaciology 49(166), 420–428. melvold, k. & hagen, j.o. 1998: evolution of a surge-type glacier in its quiescent phase: kongsvegen, spitzbergen 1964–95. journal of glaciology 44(147), 394–404. olesen, o.b. & reeh, n. 1969: preliminary report on glacier observations in nordvestfjord, east greenland. rapport grønlands geologiske under søgelse 21, 41–53. østrem, g. 1959: ice melting under a thin layer of moraine and the existence of ice cores in moraine ridges. geografiska annaler 41, 228–230. paterson, w.s.b. 1994: the physics of glaciers, 480 pp. oxford: elsevier. 76 authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: mcit@geus.dk rosa_2008:rosa-2008 01/07/09 15:49 side 76 geological survey of denmark and greenland bulletin 11, 185-204 185 faults and fractures in central west greenland: onshore expression of continental break-up and sea-floor spreading in the labrador – baffin bay sea robert w. wilson, knud erik s. klint, jeroen a.m. van gool, kenneth j.w. mccaffrey, robert e. holdsworth and james a. chalmers the complex ungava fault zone lies in the davis strait and separates failed spreading centres in the labrador sea and baffin bay. this study focuses on coastal exposures east of the fault-bound sisimiut basin, where the onshore expressions of these fault systems and the influence of pre-existing basement are examined. regional lineament studies identify five main systems: n–s, nne–ssw, ene–wsw, ese–wnw and nnw–sse. field studies reveal that strike-slip movements predominate, and are consistent with a ~nne–ssw-oriented sinistral wrench system. extensional faults trending n–s and ene–wsw (basement-parallel), and compressional faults trending e–w, were also identified. the relative ages of these fault systems have been interpreted using cross-cutting relationships and by correlation with previously identified structures. a two-phase model for fault development fits the development of both the onshore fault systems observed in this study and regional tectonic structures offshore. the conclusions from this study show that the fault patterns and sense of movement on faults onshore reflect the stress fields that govern the opening of the labrador sea – davis strait – baffin bay seaway, and that the wrench couple on the ungava transform system played a dominant role in the development of the onshore fault patterns. keywords: faults and fractures, extensional tectonics, wrench systems, sedimentary basins, basement reactivation, west greenland _______________________________________________________________________________________________________________________________________ r.w.w., k.j.w.m. & r.e.h., reactivation research group, department of earth sciences, university of durham, durham dh1 3le, uk. e-mail: robert.wilson@durham.ac.uk k.e.s.k., j.a.m.v.g. & j.a.c., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. introduction pre-existing heterogeneities in the continental crust, such as shear zones and terrain boundaries, have long been known to influence the structure and development of later deformation events (butler et al. 1997; holdsworth et al. 1997, and references therein). the sedimentary basins of the labrador sea – baffin bay region are situated west of greenland (fig. 1) and are early cenozoic failed spreading centres (chalmers & pulvertaft 2001), separated by the davis strait. the orientation of the davis strait relative to the proposed spreading centres in the labrador sea and baffin bay is consistent with the geometry of an ‘extensional transform zone’ (taylor et al. 1994). steep basement fabrics of the nagssugtoqidian orogen trend highly obliquely to these offshore structures (fig. 1) and coincide with this ‘step-over zone’ in the davis strait. fault systems fundamental to the development of sedimentary basins in the davis strait are exposed onshore in west greenland. in this project, the onshore fault systems of central west greenland were studied in order to improve © geus, 2006. geological survey of denmark and greenland bulletin 11, 185–204. available at: www.geus.dk/publications/bull 186 1000 200 km baffin island disko greenland oceanic crust transitional crust mesozoic basin unknown crust exposed shallow continental basement ? ? ? ? 50°60° 70° 65° 65° 60° 70°70° 60° 50° nuus sua q bas in iti lli f z d av is s tr ai t hi gh w est g reenland shelf n uu k ba si n c an ad a greenland nagssugtoqidian orogen fig. 2 baffin bay labrador sea border to nagssugtoqidian orogenic belt basement fabric nuussuaq nuuk inland ice aasiaat palaeogene volcanics basin hecla high ik er m iu t f z d av is st ra it sisimiut maniitsoq high nisz nssz ikermiut fz: ikermiut fault zone itilli fz: itilli fault zone nisz: nordre isortoq shear zone nssz: nordre strømfjord shear zone ubekendt ejland svartenhuk halvø u ng av a fa ul t z on e saglek basin sis im iut lin eam ent lady franklin platform sisimiut basin nuuk fig. 1. regional tectonic map of the offshore geology of the labrador sea – baffin bay area between canada and greenland. modified from chalmers & pulvertaft (2001). facing page: fig. 2. geological map of the nagssugtoqidian orogen of central west greenland (modified from escher & pulvertaft 1995). a: outline map of greenland highlighting the region covered in fig. 2b. b: geological map of the nagssugtoqidian orogen showing main lithological units and basement structures. c: topographic contour map of the central coastal area showing the field camps chosen for this study (camps 0 to 4); black dashed lines highlight major topographic escarpments. d: 3-d model view of nne-trending coastal escarpment, constructed in arcgis by draping a landsat image onto a topographic model. abbreviations used: sno, cno and nno are the southern, central and northern nagssugtoqidian orogen, respectively. itz, ikertôq thrust zone; nisz, nordre isortoq shear zone; nssz, nordre strømfjord shear zone. 187 n orth a m erican plate greenland iceland 10 km hilly relief planated relief 200–300 m > 1000 m increasing height of topography camp 0 camp 1 d 50 km disko bugt inland ice aasiaat naternaqkangaatsiaq attu sisimiut kangerlussuaq sø nd re str øm fjo rd nordre strømfjord arfersiorfik 54° 67° 69° 68° 51° n ag ss ug to qi di an o ro ge n n or th a tla nt ic c ra to n c n o n n o sn o nisz nssz itz ussu it ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ basement fabrics proterozoic archaean supracrustal rocks surficial quaternary deposits metasedimentary rocks anorthosite and ultrabasic rocks granitic intrusions orthogneisses thrusts calc-alkaline intrusions arfersiorfik and sisimiut suites granitic intrusions metasedimentary rocks archaean gneisses reworked in the palaeoproterozoic n n b c a nordre isort oq 67°10' 53° 54° camp 4 camp 2 camp 3 camp 0 camp 1 camp 4 camp 2 camp 3 0 100 200 300 400 500 600 700 800 900 1000 1100 1200 1300 1400 1500 nordre isortoq c 10 km coastal flats mountains nord re str øm fjor d n or dr e is or to q coastal escarpment greenland plate canada mapping mapping arearea a 67°20' 67°30' 67°10' 53° 67°20' 67°30' contours m above sea level mapping area nord re str øm fjo rd nord re str øm fjo rd nord re str øm fjo rd 67°40'67°40'67°40' 188 the understanding of the role played by basement reactivation in offshore basin development. in the summer of 2003 field work was carried out in an area along the coast stretching from nordre strømfjord (nassuttooq) in the north, to nordre isortoq in the south (figs 2, 3). this area was selected because it is believed that two ene-striking palaeoproterozoic shear zones, the nordre strømfjord shear zone and the nordre isortoq shear zone, were reactivated and played an important role in the development of the mesozoic to tertiary sedimentary basins. the western projection of the nordre isortoq shear zone appears to coincide with the southern faulted boundary of the sisimiut basin offshore (figs 1, 2). in addition, the offshore basins occur close to the coast in this area, and the offshore extensional faults that drop the top of the basement down to 3 km below sea level, within 10 km west of the coast, are believed to correlate with fault escarpments onshore in this area. tectonic and geological setting offshore the labrador sea and baffin bay formed during divergent plate motion between greenland and north america during the early cenozoic (chalmers & pulvertaft 2001, and references therein). the extensional basins of baffin bay and the labrador sea are separated by a bathymetric high in the davis strait (fig. 1). this transverse ridge is interpreted as a complex sinistral-shear transform fault zone, known as the ungava fault system (fig. 1; chalmers et al. 1995). extensional faulting and tectonic subsidence are thought to have commenced in the early cretaceous, at the same time as sea-floor spreading in the north atlantic south of the charlie gibbs fracture zone. opening started during the paleocene (chian & louden 1994; chalmers & laursen 1995), and sea-floor spreading appears to have ceased by the oligocene. interpretation of seismic reflection data has revealed the existence of a number of sedimentary basins offshore western greenland (chalmers et al. 1995; whittaker 1995). one such basin is the deep sisimiut basin, located in the davis strait to the west of the nordre strømfjord region (fig. 1). at about 10 km west of the coast, the top of the basement is at c. 3 km depth, while there is no cover preserved on top of basement exposed onshore. therefore, the eastern border of the basin must be a major fault, but it is located too close to the coast to have been surveyed by a seismic experiment. the orientation of this bounding fault is likely to follow the nne–ssw trend of the coastline. the western margin of the sisimiut basin is the nne–ssw-trending ikermiut fault zone (fig. 1), a transpressional flower structure developed along the transform fault (ungava fault zone) between the north american and greenland plates (fig. 2) formed during the palaeogene (i.e. it cuts early eocene strata; chalmers & pulvertaft 2001). the nukik platform lies to the south of the sisimiut basin and is separated from it by a line of eneand e-trending faults that coincide with the offshore extension of the nordre isortoq shear zone onshore (figs 1, 2). it has therefore been proposed that the faults at the southern margin of the basin developed by reactivation of basement shear zone structures in the central nagssugtoqidian orogen. these faults affect mesozoic sediments and are overstepped by paleocene sediments, so that the latest significant movement on them must have been prior to the end of the paleocene. onshore onshore exposures in central west greenland, from søndre strømfjord in the south to disko bugt in the north, comprise high-grade gneisses of the palaeoproterozoic nagssugtoqidian orogen (fig. 2; ramberg 1949; van gool et al. 2002). the nagssugtoqidian orogen is a 300 km wide belt of predominantly archaean orthogneisses, palaeoproterozoic paragneisses and intrusive rocks, that were reworked during palaeoproterozoic orogenesis (van gool et al. 2002). these basement rocks form ene-trending linear belts of steeply dipping gneisses, some of which are crustal-scale shear zones (i.e. the nordre strømfjord and nordre isortoq shear zones, fig. 2), which alternate with zones dominated by kilometre-scale fold structures (van gool et al. 2002). the nagssugtoqidian orogen is divided into three tectonic segments: the southern, central and northern nagssugtoqidian orogen (fig. 2; marker et al. 1995). the onshore research in this study lies entirely within the granulite facies orthogneisses of the central nagssugtoqidian orogen, which is bound to the north by the facing page: fig. 3. lineament map of the main study area derived from lineament mapping of a landsat tm image at 1:100 000 scale (total 1284 lineaments), using arcgis. the main rose diagram (bottom right) shows the distribution of lineaments for this map, while smaller rose diagrams (left) show results from aerial photograph analyses (at 1:20 000 scale) for each field camp. green, system 1; blue, system 2; red, system 3; yellow, system 4; purple, system 5. red box shows the position of fig. 4. 189 nordre iso rtoq nord re s trø mfjor d fig. 4 n = 295 n = 784 n = 320n = 320 n = 821n = 821 n = 295 n = 784 n = 320 n = 821 n = 1284n = 1284 n n n n 67°10´n67°10´n landsat lineaments (1:100 000)landsat lineaments (1:100 000) aerial photograph lineamentsaerial photograph lineaments (1:20 000) 000) aerial photograph lineaments (1:20 000) ese–wnwese–wnw se–nwse–nw lineaments:lineaments: 6 km6 km ne–swne–sw ene–wswene–wsw e–we–w nne–sswnne–ssw n–sn–s ese–wnw se–nw ne–sw ene–wsw e–w nne–ssw n–s n n n n camp 0camp 0 camp 1camp 1 camp 2camp 2 cam p 4 cam p 4 67°10´n nnw–ssennw–ssennw–sse lineaments: 6 km n n n n n camp 0 camp 1 camp 2 cam p 4 6 km n 190 nordre strømfjord shear zone (marker et al. 1995; van gool et al. 2002) and to the south by the ikertôq thrust zone (fig. 2). indirect topographic evidence from geomorphological investigations has suggested that late (mesozoic or cenozoic) onshore fault movements may have occurred (bonow 2004). for example, in some of the larger inlets and valleys, characteristic recent shelly marine sands can be observed up to 30 m above sea level. these are likely to have been uplifted due to isostatic rebound following glacial retreat. however, variations in elevation of these palaeoshorelines may also result from differential vertical fault movements or differential unloading. on the larger scale, the nordre strømfjord shear zone marks a major change between two landscape types (japsen et al. 2002). south of the nordre strømfjord shear zone topography is planated, with flat mountain tops forming a plateau that gradually increases in height southwards from 500 to 1000 m (fig. 2c, d). in contrast, north of the nordre strømfjord shear zone, the land has a hilly relief with a relatively flat and lowlying topography with isolated hills up to 300 m high. locally, the change in landscape type occurs across a more than 500 m high, ene–wsw-oriented escarpment that drops down to the north (fig. 2). there is also a pronounced nne–ssw-oriented escarpment almost 1 km high that drops down to the west between nordre strømfjord and nordre isortoq (fig. 2). this major escarpment separates low-lying (50–150 m high) coastal flats to the west from the much higher (500 m+) mountains to the east (fig. 2) and can be traced for over 80 km, from sisimiut in the south to nordre strømfjord in the north. similar escarpments can be observed in the near offshore on both bathymetric and horizon maps of depth to basement, thus supporting the theory that onshore structures reflect those offshore. methods the present study combines regional to outcrop-scale mapping and regional studies of remotely sensed data to determine fault-fracture geometries, distribution, relative timing and kinematics in selected key areas of central west greenland. regional studies comprised satellite image and aerial photograph analysis at a variety of scales (1:500 000; 1:100 000; 1:20 000) in order to identify lineaments and other geological structures (e.g. variations in lithology, fabric intensity, faults, fractures). field investigations were carried out in the well-exposed precambrian basement rocks in key areas of interest that were identified during aerial photograph analysis prior to departure. a number of field camps were used during the mapping (fig. 2c). camp 0 was located at inussuk, a site visited previously in 2002 (japsen et al. 2002). camps 1 and 2 were located on the north and south shores of nordre strømfjord, while camp was on the north shore of nordre isortoq. camp 3 was farther inland, and is not analysed further in this study. during field work, fault and fracture systems were mapped, and the following structural data were collected for statistical/structural analysis: • fault attributes including: orientation; kinematics; fault surface characteristics; mineralisation. • relative age relationships. • structural/statistical analyses to determine kinematic patterns. over 200 pseudotachylite and mica-bearing fault-rock samples were also collected from different fault sets at various localities in order to date the fault movements using 40ar/39ar geochronology (results to be discussed elsewhere). all field data were geospatially located (to 5 m resolution) using global positioning system (gps) waypoint collection, and were subsequently stored in a computer database with links to geographic information system (gis) based maps. fault and fracture characterisation fractures include all brittle structures such as joints, fissures, cracks, veins, etc. that are not faults, bedding or cleavage surfaces, and are larger than the grain size of the rock. in general, fractures are defined as dominantly tensile (mode i) cracks, and as such, they are associated with characteristic stress, strain and displacement fields. they are distinguished from small faults by distinctive surface textures and lack of shear displacements. faults are mapped where distinct offsets have been identified, often with a development of slip striae on the surface (slickenlines). criteria for determining the sense of movement were based on methods outlined in mcclay (1987) and petit (1987). in the present study faults were classified as: • normal (extensional dip / oblique-slip fault). • reverse (compressional dip / oblique-slip fault). • strike-slip faults (dextral or sinistral). the classification of faults and fractures into systems was primarily based on the orientation of the structures (i.e. 191 trend of lineament or strike of plane). structures with different orientations can reflect different deformation phases, but in complex fault zones developed in three-dimensional (3-d) strain fields, multiple fault and fracture orientations may develop during a single event (see e.g. de paola et al. 2005). therefore further classification needs to be applied, either through systematic fracture properties (such as surface type or mineralisation), or through kinematic studies, in order to determine if only one or several phases of deformation are apparent. fault and fracture measurement technique in order to accurately classify the fault and fracture systems, populations of at least 50 fractures/faults were measured at most localities. faults were classified according to type (normal/reverse dip-slip faults or dextral/sinistral strike-slip faults). fault orientations and the directions of slickenlines (when observed) were measured, and the following characteristics recorded: • surface shape: the overall fracture shape (metre-scale) was described as listric, planar, undulating or irregular. • surface roughness character: the surface roughness character (millimetre scale) was described as smooth, rough or slickenside (striae). • other features: some fractures/faults have a filling of iron oxide precipitates, quartz crystals, epidote, or preferential growth of other crystals on the surface showing the slip direction. special types of fractures such as conjugate shear fractures, en échelon fractures, plumose jointing etc. were noted if present. = + + + or system 1 system 4 (yellow) system 2 system 3 youngestoldest 1 km n fig. 4. age relationships interpreted from cross-cutting relationships of lineament systems derived from aerial photographs for camp 1. four dominant lineament trends are apparent: n–s (system 2), nne–ssw (system 4), ene–wsw (system 1), and nnw–sse (system 3). through crosscutting relationships a relative order of fault development is apparent, as indicated across the bottom of the image. 192 lineament mapping lineament maps for the central nagssugtoqidian orogen (fig. 3) were plotted from landsat tm images and aerial photographs at a variety of scales (1:500 000 and 1:100 000 for landsat images and 1:20 000 for aerial photographs). images were georeferenced and displayed in a gis environment and the lineaments picked by hand. after interpretation, lineaments were then refined using digital terrain model (dtm) analysis and compared to pre-existing geological maps of the region (e.g. henriksen et al. 2000). as the data are stored in a gis, attribute data for each lineament (i.e. trend; length; offset; other features) were also measured or calculated and stored. spatial analysis and rose plotting tools in arcview gis were used to analyse the orientation (fig. 3) and spatial distribution of these structures. a more detailed analysis of selected areas was then carried out at 1:20 000 scale using aerial photographs. as image resolutions are much higher in aerial photographs (2 m pixel size), particular attention was paid to how the lineaments interact with topography (e.g. v-ing into valleys, etc.) to gain a better understanding of their overall geometry. generally all lineaments picked appear to have a steep dip as only minor interactions with topography were observed. attention was also paid to cross-cutting relationships between lineaments in an attempt to determine the relative timing of structures (fig. 4). lineament systems in total 1284 lineaments have been identified from landsat tm images (pixel size 30 m) at 1:100 000 scale (fig. 3). lineaments derived from both landsat and aerial photographs have been grouped into systems based on their orientation. five main lineament systems (n–s, nne– ssw, ene–wsw, ese–wnw, and nnw–sse) have been identified (fig. 3; table 1). system 1 structures (green; figs 3, 4) are oriented ene– wsw (~060–090° trend), and are pervasively distributed across the region. this system has a trend similar to the nordre strømfjord and nordre isortoq fjords, and lies parallel to the regional basement fabric (foliation, gneissic banding, and shear zones; van gool et al. 2002). note that as these lineaments may represent either faults or basement fabrics, care must be taken when analysing these quantitatively. in an attempt to minimise the amount of oversampling, only the most pronounced lineaments (e.g. most weathered out) that mark a distinct change in structure were mapped, while those that are clearly basement fabrics (i.e. those showing ductile features such as folds) were not. system 2 lineaments (blue; figs 3, 4) are n–s oriented (trend ~350–010°), and often show sinistral offsets of preexisting structures (basement fabrics). this system can be traced from nordre isortoq to the northernmost part of the investigated area (fig. 3), and previous investigations indicated that they may continue as far north as aasiaat, disko and nuussuaq (japsen et al. 2002). the fault zones are closely spaced (100–500 m), and strike-slip separations of up to 30 m have been observed. system 3 lineaments (red; figs 3, 4) are nnw–sse oriented (trend ~140–170°). they are closely spaced (50– 100 m), and offsets of marble beds show net dextral separations in the order of 20–40 m (fig. 4). this system is most pronounced in the nordre strømfjord shear zone, and less dominant in the nordre isortoq shear zone (see rose diagrams in fig. 3). system 4 lineaments (yellow; figs 3, 4) are oriented nne– ssw (trend ~010–040°) and are strongly developed in the nordre strømfjord shear zone region (camps 0 and 1, fig. 3). these structures show apparent sinistral strikeslip separations of up to 400 m in the westernmost part of the study area. the spacing between them increases from approximately 500 m at the coast, to approximately 2 km farther inland. the same lineament directions were encountered at camps 2 and 4, south of nordre strømfjord, where these structures are shorter and more discontinuous, possibly due to differences in rock type and fabric between the two areas. system 5 lineaments (purple; fig. 3) consist of e–w to ese–wnw (trend ~090–120°) -oriented structures. the valleys that distinguish this system are generally 10–30 m wide and have a curved trend. this system appears to be mostly localised into two specific areas: the first of these lies in the fold belt south of nordre strømfjord (fig. 3), and the second is located south of sisimiut (fig. 1). relative timings figure 4 shows an aerial photograph of an area around camp 1 where an apparent order of lineament development can be deduced. the oldest structures appear to be system 1 (green), and in this area these structures appear to be basement fabrics in the form of alternating layers of semipelite and marble up to 100 m thick (henriksen et al. 2000). system 2 (blue) structures show sinistral offsets of these lithological layers, while system 3 structures (red) show dextral displacements. in fig. 4, system 3 structures appear to dominantly cross-cut system 2 structures, but this is not always the case as in some areas the reverse is true (system 2 cross-cutting/displacing system 3). as there 193 is evidence for systems 2 and 3 mutually cross-cutting each other, and also because apparent movements are compatible with a conjugate system of strike-slip faults, it is possible that they are contemporaneous. cross-cutting all other systems are the nne-trending system 4 lineaments, suggesting that they are likely to be the youngest structures, or at least have experienced the most recent movements. other areas show a similar pattern of events, although some system 1 structures show evidence for younger movements (reactivation?), especially in areas around camps 2 and 4. system 5 is not represented in fig. 4 as it was not observed at camp 1. this system is marked by quite wide (30–50 m) valleys, thus making its displacements difficult to determine; however, as it is quite pronounced, it may be a more recent system. field observations four key areas were chosen for detailed fracture and fault analysis in the field (fig. 3), based on their structural interest (i.e. their potential to enable all systems to be analysed) and accessibility. the first objective of the field work was identification of the lineaments picked from the aerial photographs. in most cases field observations proved that the lineaments correspond to major fault structures, many of which are weathered out to leave gorges and river valleys (fig. 5a–c). however, whilst many of the ene– wsw-oriented system 1 structures are faults, others also correspond to basement fabric features, such as strongly foliated zones, lithological contacts and shear zones (fig. 5d). therefore care must be taken in any quantitative geometric or spatial analysis of this system. after a regional reconnaissance from each field camp, detailed structural analysis was carried out. ninety outcrop locations were investigated in the four camps, distributed along the coast between nordre isortoq and just north of nordre strømfjord (fig. 2). in total c. 1700 faults and fractures were measured and described. fault geometries a wide range of fracture orientations were observed (fig. 6a), with dominantly n–s and nnw–sse strikes and an overall mean fracture plane of 167/89e. various fault orientations can be separated out in the field (dominant trends are n–s and ene–wsw), showing a range of slip movements and shear senses (fig. 6b–f). dominant fault movements appear to be strike-slip (71% of faults recorded show strike-slip movements), although extensional and compressional faults were also apparent. • basement-parallel to subparallel • multiple phases of movement • closely spaced (100–500 m) • displacements range between 0.2 and 30 m for individual faults • may show an en échelon to irregular trend • dominant fracture/joint trend is associated with this system • closely spaced (50–100 m) • marble layers show displacements in the order of 20–40 m • major subvertical faults and fault zones • generally associated with wide (> 50 m) valleys • exposed fault cores show complex fracture sets associated with strike-slip movements • localised to the fold belt south of nordre strømfjord and north of the nordre isortoq shear zone • prominent structures at regional scale (i.e. from landsat and aerial photos) but not at outcrop • spatially associated with compressional faults (i.e. system 1 reverse faults) • normal (dip-slip) • reverse (dextral oblique-slip) • dextral and sinistral strike-slip • sinistral strike-slip • normal (dip-slip) • dextral strike-slip • normal (dip-slip and oblique-slip) • sinistral strike-slip • dextral strike-slip system 1 system 2 system 3 system 4 system 5 ene–wsw n–s nnw–sse nne–ssw e–w to ese–wnw comments sense of movementlineament system orientation table 1. main characteristics for each fault system, identified from remote sensing and outcrop studies 194 a set of ene–wsw-oriented faults (green planes/mean poles in fig. 6) appear to reactivate strong basement fabrics in the nordre strømfjord and nordre isortoq shear zones. these faults correspond to system 1 lineaments and exhibit various forms of fault movement (e.g. extensional, compressional and strike-slip; fig. 6b–f). faults corresponding to systems 2 (n–s, blue), 3 (nnw–sse, red) and 4 (nne–ssw, yellow) can also be easily distinguished from the fault data in fig. 6. however, lineament system 5 (e–w to ese–wnw, purple) is not apparent. as previously discussed this system appears to be a more geographically localised system (i.e. local to areas south of the nordre strømfjord and nordre isortoq shear zones), and correspond to zones dominated by reverse fault movements (fig. 6b) and a small number of ese-trending dextral strike-slip faults (fig. 6e). fault systems corresponding to lineament systems 2, 3 and 4 appear to consist of parallel fault zones separated by non-faulted, but generally strongly fractured rock. the fault zones range typically between 1 and 50 m in width (fig. 5b) and consist of multiple parallel faults with variable spacing. these zones are commonly located in pronounced valleys and gorges (fig. 5a), so characterisation of fault planes was often difficult as the valley floors are generally covered by recent sediment and vegetation. bb lunate fracture indicatinglunate fracture indicating dextral strike-slip movements dextral strike-slip movements rm structuresrm structures indicating dextralindicating dextral strike-slip strike-slip ww e a cc raised beachraised beach d system 1 system 1 system 4system 4 sinistral strike-slip movementssinistral strike-slip movements (inferred from ramp-step geometry) (inferred from ramp-step geometry) steeply dippingsteeply dipping marble unit marble unit nnennesswssw raised beach system 1 system 4 steeply dipping marble unit nne n s ssw nnessw ew a c d b lunate fracture indicating dextral strike-slip movements sinistral strike-slip movements (inferred from ramp-step geometry) rm structures indicating dextral strike-slip fig. 5. field identification of lineaments picked from landsat and aerial photographs. a: two major gorges/valleys trending ene and nne near camp 1; these correspond to major lineament systems 1 and 4, respectively. fault core exposed within the nne-trending stream bed (system 4) shows evidence for sinistral strike-slip faulting (fig. 5b); fault movements on the ene-trending basement-parallel valley (system 1) were not identified. b: photograph of subhorizontal, sinistral strike-slip slickenlines observed within the fault core of the nne-trending fault (system 4) identified in fig. 5a. c: some ene-trending (system 1) lineaments correspond to basement fabrics such as steeply dipping (and tightly folded) marble units. d: other basement parallel lineaments, however, do show evidence for brittle fault movement, as identified in this ene-trending fault core (fault movement criteria defined by secondary fracture indicators, i.e. rm and lunate fractures, outlined in petit 1987). 195 fault kinematics in addition to the characterisation of the faults and fractures in terms of their trend and distribution, they can also be described according to their movement patterns (see table 1). the nature and timing of tectonic events that are responsible for the formation of these fault-fracture systems is quite complex. multiple directions of slickenlines on several fault surfaces indicate that many faults were either reactivated or that individual faults exhibit curved movement trajectories consistent with complex strain histories. strike-slip faulting strike-slip slickenlines account for 71% of those observed and were observed on all main fault geometries or systems. multiple orientations of strike-slip faulting are common in wrench-dominated fault systems due to the development of riedel, p and x shears (e.g. woodcock & schubert 1994). basement parallel faults (system 1, green) show both dextral and sinistral movements (note, rm structures – petit 1987 – associated with r-shears suggest dextral movements on basement faults in fig. 5d). nnw-trending (system 3, red) faults appear to correspond to dextral movements. n-trending faults (system 2, blue) appear to poles to planes slickenlines a b c fed n n n nnn fractures n = 983 mean fracture orientation reverse faults n = 80 normal faults n = 141 strike-slip faults, n = 339 (shear sense uncertain or unassigned) dextral faults n = 96 sinistral faults n = 107 e +2s +4s +6s +8s +10s +12s +14s +16s +18s +20s +22s +24s e +2s +4s +6s +8s +10s +12s +14s e +2s +4s +6s e +2s +4s +6s +8s 10s e +2s +4s +6s +8s e +2s +4s +6s fig. 6. lower hemisphere, equal area stereographic projections of all fault and fracture data collected at camps 1, 2 and 4 (total number of measurements = 1746). a: fractures (i.e. planes showing no evidence for slip). b: reverse faults. c: normal faults. d: strike-slip faults with undetermined sense of movement. e: dextral strike-slip faults. f: sinistral strike-slip faults. black dots, poles to planes of fault and fracture surfaces. red dots, slickenlines. mean fault planes are also shown, coloured according to lineament/fault systems identified in fig. 3. poles to fault and fracture surfaces are contoured using a gaussian weighting function; n, number of measurements for each plot. in the labels, e corresponds to the background value (calculated as number of points/100), while s = standard deviations above this value. 196 aa bbb ccc aa 1st order fractures normal faults associated slickenlines 1st order fractures 2nd order fractures strike-slip faults (sinistral) wsw ene 2 m2 m2 m 0.5 m0.5 m pseudotachylite en échelon footwall joints (subvertical dip) a 0.5 m nd n = 56 fig. 7. a: panoramic photograph showing exposures of a series of parallel ene-dipping extensional faults, in the vicinity of camp 1. b and c: close-up photographs showing en échelon fracturing on the footwall of normal faults and pseudotachylite fault exposure in more detail. d: lower hemisphere, equal area stereographic projection of poles to planes, and associated slickenlines, for faults and fractures observed at the outcrop of fig. 7a. fault orientations and kinematics suggest ene–wsw extension as indicated by stress arrows (red); n, number of measurements. 197 show both dextral and sinistral movements, while nnetrending faults (system 4, yellow) are sinistral structures (fig. 6d–f). extensional faulting although all five fault systems show dominantly strikeslip movements, normal and oblique-slip components of displacement were also recorded on some sets. these faults appear to have two dominant orientations, n–s (system 2) and ene–wsw (system 1; fig. 6c). in some areas nnw–sse (system 3) -oriented structures also appear to be normal faults (fig. 7), but these are not the dominant orientations in bulk analyses (fig. 6). field observations suggest that strike-slip movements post-date dip-slip. compressional faulting reverse faults (figs 6b, 8) appear to be confined to areas close to camps 2 and 4, and to be spatially associated with system 5 lineaments. these faults strike parallel or subparallel to basement structures (ene–wsw to e–w, fig. 6b) and exhibit dextral-oblique slickenlines, which plunge towards the ese (figs 6b, 8). as these compressional or thrust faults are only found on the southern shore of nordre strømfjord and the northern shore of nordre isortoq (i.e. abutting against major basement shear zones) it is possible that these structures are the result of local transpressional thrust faulting, which may be linked to steps in the en échelon sinistral fault system (system 4). field observations suggest these compressional faults post-date most other fault and fractures. however, at camp 4, a n–s sinistral strike-slip fault appears to cross-cut these thrusts (fig. 8b). reverse faults sinistral strike-slip dextral strike-slip strike-slip fault (movement undetermined) n = 91 n reverse faults basement-parallel dextral faults mean slickenlines orientation on reverse faults 1 m 1 m ene ene aa c wsw wsw 48-120 078 / 50 s 078 / 50 s 1 m n s wsw b a 078 / 50 s 078 / 50 s 48-120 48-120 c ene fig. 8. a: exposure of a localised set of reverse faults near camp 4. b: lower hemisphere, equal area stereographic projections of poles to structures observed at the locality of fig. 8a. three dominant fault sets are apparent: basement parallel – i.e. ene-trending – reverse and dextral strike-slip faults, and a set of sinistral faults oriented n–s (e.g. along the large rock face in shadow). slickenline orientations and relative fault movements suggest ese–wnw compression as indicated by stress arrows (red); n, number of measurements. c: photograph of surface of reverse fault, showing dextral-oblique slickenlines (mean slickenline orientation 45/120, see stereonet). note that faults coloured in red here highlight thrust faults and do not refer to system 3 faults as in other figures. 198 fractures and joints a diverse array of fracture orientations was recorded. dominant orientations vary from nw–se through to nne– ssw. the orientation of the mean plane is nnw–sse. all fractures recorded showed no evidence for shear movement (i.e. rough surfaces and with no apparent offsets) and are thus interpreted as opening mode 1 fractures and suggesting extension directions varied from c. e–w to ne– sw. interpretation and discussion an overall summary of each of the fault systems identified through remote sensing (i.e. lineament mapping) and field studies is presented in table 1. in this section we discuss the possible interpretations and implications of these observations. fault development system 1 (ene–wsw) faults and fractures appear to be the oldest structures, however multiple slip vectors and apparent fault movements suggest that there has been activity on this system during later tectonic episodes (note that system 1 faults are apparent in all stereoplots for all fault types, fig. 6b–f). these lie parallel to the pre-existing nagssugtoqidian basement fabric, which dates at c. 1.8 ga (van gool et al. 2002). cross-cutting relationships interpreted from analysis of aerial photographs suggest that the next systems to develop were systems 2 and 3 (fig. 5). it is difficult to determine if one of these systems predates the other as mutually cross-cutting relationships can be seen; however, it does appear that system 3 is the more pervasive system and thus may be more recent. strike-slip movements and offsets on systems 2 (n–s) and 3 (nnw–sse) suggest that, if active at the same time, these would represent a strike-slip conjugate system. in such cases the inferred extension vector would trend ene– wsw, subparallel to system 1 foliation-parallel faults. this extension vector is also consistent with the dip-slip fault movements seen locally on these same fault systems. these strike-slip movements appear to be preceded by dip-slip extensional movements. system 2 (n–s) is the dominant extensional fault orientation in the area (fig. 6c), while some localities showed small populations where nnw–sse-oriented extensional faults represent the preferred trend (e.g. fig. 7). these faults are indicative of e– w to ene–wsw extension. this extension cannot, however, explain the apparent basement-parallel (system 1) extensional faults, which suggest an apparent nnw–sse extension. these ene-trending normal faults have also been observed in seismic interpretations and are thus important structures regionally. these two extensional fault sets show a quadrimodal fault distribution, i.e. four sets of fault planes (fig. 6c). if regarded as two separate fault sets, this geometry would suggest two separate extension directions (e–w and nnw–sse); however an alternative to this is that these faults formed contemporaneously under three-dimensional strain (reches 1983; nieto-samaniego & alaniz-alvarez 1997). as one set of normal faults trends parallel to a preexisting plane of weakness (e.g. basement fabric) it is likely that basement reactivation played a role in the development of these faults, and that this has lead to formation of extensional faults oblique to the regional extension. the regional tectonic setting would fit with 3-d strain as the area borders the transfer zone between two extensional basins, i.e. is a transtensional deformation zone (dewey 2002; de paola et al. 2005). geoffroy et al. (1998) recorded similar fault geometries and kinematics farther north in disko and nuussuaq (fig. 1). their interpretation is that strike-slip and dip-slip faulting developed during a single tectonic episode of wsw– ene extension, which is in agreement with a model of 3-d transtensional strain. as the stereoplots in fig. 6 show, strike-slip faults are the dominant fault type in the area (as mentioned above, 71% of all faults measures are strike-slip) and these appear to post-date extensional movements. all fault systems show evidence for strike-slip movements. system 1 (ene–wsw) exhibits both dextral and sinistral senses of shear, systems 3 (nnw–sse) and 5 (ese–wnw) show dextral shear, while systems 2 (n–s) and 4 (nne–ssw) are dominated by sinistral shear movements. system 4 faults (nne–ssw) appear to cross-cut all other fault sets, and are characterised by major fault zones (fig. 5). these major sinistral strike-slip structures lie subparallel to the sinistral ikermiut and ungava fault zones that dominate the davis strait offshore (fig. 1). assuming a ~nnetrending sinistral wrench system for the study area, strikeslip fault movements on each system appear to correlate closely with synthetic (r) and antithetic (r’) reidel shears, and also with synthetic p and antithetic x shears typical of a plane strain wrench tectonic regime (fig. 9; woodcock & schubert 1994). compressional faults appear to be relatively late structures (although cross-cut by ~n–s-trending sinistral faults) and are localised in areas of strong basement fabric (i.e. 199 shear zones). these faults strike parallel to basement fabrics and indicate an oblique compression (from the ese or se; fig. 8). offshore there is evidence for thrusting in a similar orientation along the ikermiut fault zone (fig. 1). positive flower structures have been identified (chalmers & pulvertaft 2001) and are interpreted as inversion structures formed at a restraining bend during sinistral strike-slip along the ungava transform fault, during the early eocene (c. 54–49 ma, chalmers & pulvertaft 2001). if trends of basement shear zones (outlined in fig. 2) are continued along strike offshore they appear to coincide with these transpression zones within the ikermiut fault zone. it is possible that thrusts observed onshore have formed in a similar way to those offshore with basement shear zones acting as restraining bend structures, thus leading to localised compressional zones. furthermore, slickenlines on the reverse faults suggest a compression from the ese or se (figs 6b, 8), which is consistent with the compressional axis for a sinistral wrench system (i.e. ne–sw extension and nw–se compression; fig. 9). regional comparison and implications a key prerequisite for building tectono-stratigraphic models is being able to date each tectonic event. as all onshore exposures in this part of west greenland are in precambrian basement rocks, there are no stratigraphic markers for constraining the timing of phanerozoic tectonic events. relative timing has been inferred from various cross-cutting relationships in the field and from lineament analysis, but is open to different interpretations. in the absence of age data that constrain the absolute age(s) of fault activity, comparisons with offshore models and with data collected in other onshore areas are used here to infer ages for events in our tectonic model (see table 2 for a summary). regional onshore correlations farther north in the region of disko and nuussuaq, onshore faulting episodes can be dated relative to the deposition of basaltic lavas and the sedimentary systems during campanian to eocene times (geoffroy et al. 1998; storey et al. 1998; chalmers et al. 1999; dam et al. 2000). dam & sønderholm (1998), dam et al. (2000) and dam (2002) document at least three phases of faulting recorded in the sedimentary record prior to paleocene volcanism. cretaceous–paleocene sediments on nuussuaq show distinctunconformities, with incised valleys and submarine canyons, reflecting disturbances in early campanian, maastrichtian and early paleocene times. these unconformities and channels are thought to have formed in response to structural movements associated with regional ne–sw rifting (and also the arrival of the north atlantic plume in the latter two cases). 40ar/39ar dating has revealed that volcanism commenced in west greenland between 60.9 and 61.3 ma and that 80% of the paleocene lava pile was erupted in less than 1 ma (storey et al. 1998). these lavas show a distinct coastal flexure (geoffroy et al. 1998, 2001; larsen & pulvertaft 2000), presently expressed by seaward dipping basalt lavas. this flexure has an arcuate course, striking nw–se in southern svartenhuk halvø and northern ubekendt ejland, turning through n–s in south-west ubekendt ejland to ne–sw in north-west nuussuaq and finally to n–s in north-west disko (fig. 1; geoffroy et al. 1998, 2001; larsen & pulvertaft 2000). numerous dykes cut these lavas (e.g. see figs 4 and 5 in larsen & pulvertaft 2000), and so do various fault sets (geoffroy et al. 1998). the timing of the various phases of volcanic eruption, dyke emplacement, and block faulting relative to one another is still a matter of debate. geoffroy et al. (1998) presented detailed structural evidence suggesting that fault and dyke intrusion on disko took place during tilting of lava sysr` r p x x r p r` a b nne–ssw (10–15°) wrench system: 55–60° extension, 145–150°compression sy st em 4 system 1 system 5 system 3 system 2 nfig. 9. a: diagram showing fault systems and their corresponding movements. b: strain ellipse for a nne–ssw (~010–190°) -oriented sinistral wrench system, showing riedel (r and r’), p and x shears (woodcock & schubert 1994). also shown are the regional stress vectors (σσσσσ1 and σσσσσ3). systems 2 (sinistral) and 5 (dextral) correspond to r’ and r shears, while systems 1 (dextral) and 4 (sinistral) correspond to p and x shears. system 3 corresponds to normal fault sets in fig. 9b; however, dominant movements on this system were dextral. 200 tems. geoffroy et al. (2001) then stated that nw–seoriented, flexure-parallel dykes in southern svartenhuk yield dates of around 54.6 ± 0.6 ma. this suggests that most of the coastal flexure is of eocene age (or later?). further evidence for this comes from north-west nuussuaq where tilted lavas (larsen & pulvertaft 2000) have been dated at ~53 ma (storey et al. 1998). systems 2 (n–s) and 3 (nnw–sse) extensional faults in this study appear comparable to faults that cross-cut these paleocene basalt lavas (geoffroy et al. 1998). taking the dates outlined above for dyke emplacement which is believed to be associated with faulting, it would appear that our system 2 and 3 faults were active during eocene times. taking all these onshore tectonic timings into account it would then appear that the faults observed may have been active from late cretaceous (maastrichtian) through to eocene times (dam & sønderholm 1998; geoffroy et al. 1998, 2001; chalmers et al. 1999; dam et al. 2000; larsen & pulvertaft 2000; dam 2002). normal fault orientations similar to systems 2 (n–s) and 3 (nnw–sse) occur in and around the nuussuaq basin (fig. 1; geoffroy et al. 1998; chalmers et al. 1999) and are consistent with either ene–wsw (geoffroy et al. 1998) or e–w (chalmers et al. 1999) extension. chalmers et al. (1999) proposed that the n–s faults formed by e–w-oriented crustal extension, while associated ese–wnw faults formed as a consequence of reactivation of shear zones in the underlying basement. normal faults are dominantly n–s onshore (fig. 6c), fitting with this model proposed by chalmers et al. (1999). however, this implies that while baffin bay in the north and labrador sea in the south were undergoing ene–wsw extension (deduced from dominant fault trends and earliest magnetic anomaly trends, chalmers & pulvertaft 2001), southern west greenland and the nuussuaq basin were undergoing e–w extension. a better explanation is that the davis strait at this time (i.e. prior to the onset of seafloor spreading) lay in a transfer/step-over zone between two extensional basins, and that it was strongly influenced by basement fabrics such that this region experienced complex 3-d strain associated with regional ene–wsw extension (fig. 10). onshore normal fault sets form a quadrimodal fault distribution (four sets of fault planes, fig. 6c) consistent with 3-d strain. this faulting is then subsequently dissected by n–s (system 2) and nne–ssw (system 4) -oriented faults during the eocene (chalmers et al. 1999). the itilli fault zone (fig. 1) is one such nne–ssw-oriented structure cutting through north-west nuussuaq. this fault zone appears to be a left-lateral splay from the northern extension of the ungava fault zone in the davis strait (chalmers et al. 1999). • uplift? • possible reactivation of systems 1–4 as normal faults • faults consistent with nne-oriented sinistral wrench system: nne-trending (system 4) faults and basement-parallel (system 1) faults active as antithetic x and synthetic p shears • systems 2 (reactivation) and 5 active as reidel shears • local transpressional thrust faulting observed near camps 2 and 4, associated with steep basement fabrics • systems 1, 2 (and 3?) faults all active as extensional faults during ne–sw to ene–wsw extension (3d strain)? • uncertain? • possible system 1 ene–wsw foliation-parallel faulting prior to late mesozoic? • subsidence • nto nne-trending sinistral transverse fault system (ungava fault zone). • local transpressional and transtensional faulting (e.g. ikermiut fault zone) • n-trending faults in davis strait (e–w extension) leading to formation of the sisimiut basin offshore • nw-trending normal faults in the labrador sea (sw–ne extension) • wsw–ene faults bordering the sisimiut basin to the south were active offshore during events 2 and/or 3 • uncertain? 5 (youngest) 4 3 2 1 (oldest) pliocene to pleistocene tilting eocene labrador sea-floor spreading (ungava system) late cretaceous to early paleocene extension early or middle to late cretaceous extension and thermal subsidence proterozoic and later localised reactivation onshore tectonic structures offshore tectonic structuresevent # timing and event table 2. proposed event stratigraphic model and apparent correlation with offshore events 201 fig. 10. proposed two-stage tectonic model for the tectonic evolution of upper mesozoic – cenozoic extension within the nagssugtoqidian orogen. stage 1(a): nand ene-trending normal faults and dextral basement reactivation due to ene–wsw extension. stage 2 (b): nand nne-trending sinistral strike-slip faulting, and associated strike-slip wrench tectonic systems, with compressional structures (reverse faults) forming in zones of basement anisotropy (e.g. shear zones). block diagrams show schematic cartoons outlining fault patterns observed onshore, while maps show the regional context, based on correlations between onshore and offshore fault structures. labrador sea baffin bay possible oceanic crust ikermiut fault zone (pop-up flower structure) n 200 km b nne-trending sinistral wrench system labrador sea ba se m en t a nis otro py ba se men t a niso tro py base ment an isotropy base ment an isotropy ba se m en t a ni so tr op y n nordre strømfjord nordre isortoq sisimiutesisimiute basinbasin 3d strain zone: quadrimodal fault patterns (strain partitioning?) n 200 km u ng av a f z u ng av a f z 2d strain zone: simple andersonian faulting baffin bay ? nagssugtoqidian orogen d av is st ra it stage 1: late cretaceous–paleocene extension (ne–sw to ene–ssw). pre-existing basement anisotropy (nagssugtoqidian orogenic belt) appears to influence fault patterns in the davis strait. possible strain partitioning between basementparallel and n–s-trending normal faults. stage 2: nne-trending eocene sinistral wrench system (ne to ene extension, se to sse compression) reactivating earlier extensional faults and basement-parallel structures (e.g. thrust faults at camp 4, and the ikermiut fault zone). n nordre strømfjord nordre isortoq a b a n–s normal faults basementparallel faults fractures 202 correlation with fault structures offshore fault patterns offshore in the davis strait exhibit similar c. n–s and ene–wsw dominant orientations. the south and east bounding faults to the sisimiut basin strike ene– wsw and n–s respectively (fig. 1). significantly, the southern margin of the sisimiut basin is coincident both in orientation and location with a major basement shear zone (the nordre isortoq shear zone), and is likely to have exerted a similar structural control to that interpreted for onshore. block faulting has been dated via drilling as having taken place between the late campanian and late paleocene (christiansen et al. 2001; dalhoff et al. 2003), indicating that these extensional faults are of similar age to those associated with valley incision on nuussuaq and north disko (dam & sønderholm 1998; dam et al. 2000; dam 2002). the ungava fault zone with its associated fault systems (e.g. the ikermiut fault zone) is the most prominent structure in the davis strait (chalmers & pulvertaft 2001). this nne–ssw-oriented structure is interpreted as a transform fault zone showing sinistral shear, and has been linked to the itilli fault zone (fig. 1; chalmers et al. 1999). system 4 (nne–ssw) faulting onshore, around nordre strømfjord, is consistent with late sinistral strike-slip movements, and it is reasonable to suggest that this system is of similar age. as already discussed, offshore evidence for sinistral strike-slip movements can be seen in the ikermiut fault zone on the western margin of the sisimiut basin (chalmers & pulvertaft 2001) where transpressional thrusts (similar to compressional flower structures modelled in dooley et al. 1999) appear to have formed in the restraining bend of a strike-slip fault (see fig. 6 of chalmers & pulvertaft 2001). these thrusts cut early eocene mudstones, but are overlain by late eocene sediments, providing further evidence for timing of these movements. this sinistral shear is thus a consequence of left-lateral movement of the canadian plate relative to the greenland plate along the ungava transform system during sea-floor spreading in the labrador sea (fig. 1). evidence for neotectonic faulting? chalmers (2000) presented evidence for neogene uplift in offshore areas of central west greenland, while recent onshore topographic and apatite fission-track data analysis has identified similar neogene activity (japsen et al. 2002, 2005). a common observation in this field area is the presence of raised beaches and palaeoshorelines, up to elevations 30 m above present sea levels (fig. 5a). they are probably the result of isostatic readjustment following the removal of pleistocene ice load. topographically, the region appears to be divided into blocks, split by enetrending fjords and nne-trending escarpments (fig. 2c). these must be quite recent features as they have not been eroded during glacial activity (and may in fact be a consequence of it). the trend of system 4 faults (and also locally those of system 2) is generally parallel to the pronounced nne-trending escarpment from nordre isortoq to nordre strømfjord onshore, and also to a similarly trending scarp near offshore (identified in bathymetry maps), and it is possible that these faults have been reactivated as normal faults during a recent tectonic event. this conjecture still needs to be verified, as the main escarpments were not studied in detail during our field work. summary the observed fault and fracture systems reflect a brittle tectonic history that is ultimately related to far-field plate movements, uplift and basin formation. the development of mesozoic to cenozoic basins offshore west greenland appears to be strongly controlled by faults. therefore, knowledge of the fracture systems in the exposed precambrian basement provides a valuable insight into fault geometries and kinematics during the development of offshore basins and potential hydrocarbon reservoirs. it also provides insights into the possible influence of basement reactivation. several possible tectonic-event models may be constructed for this region given the lack of definite ages for structures. table 2 shows a basic summary of the relative timings of fault systems identified in this study relative to regional offshore tectonic models, while fig. 10 presents a model for fault development based on the observations and correlations made in this study. the absolute timing of the fault activity onshore, as deduced from correlation to other fault systems with known ages, needs to be tested by dating of fault rock samples. a simple two-stage model has been outlined to explain the complex fault patterns exhibited in onshore exposures of the central nagssugtoqidian orogen (fig. 10). the brittle tectonic evolution of the region appears to be dominated by ne–sw extension, which is consistent with the opening of the labrador sea and baffin bay. only slight variations in the regional stress field are required to account for the diversity of fault orientations. according to chalmers & pulvertaft (2001) there was a 15° counter 203 clockwise rotation in spreading direction between the paleocene and the eocene in the labrador sea as opening started between greenland and europe, which is consistent with the two-stage model outlined in fig. 10. in the early stages of opening, faulting was dominated by extensional structures (under 3-d strain conditions), favouring an e– w to ene–wsw extension (fig. 10a); however, as the ungava transform fault developed, faulting became more wrench dominated (2-d plane strain), and suggests ne– sw extension (fig. 10b). variations in fault geometry reflect these changes in the regional stress field. however, the influence of basement structure also appears to have played an important role throughout (e.g. extensional faults not normal to the extension direction, and the apparent localised compressional zones associated with intense basement fabrics). although most faults observed onshore trend highly obliquely to basement fabrics, fault patterns do appear to vary in areas of intense pre-existing structure (such as the nordre strømfjord and nordre isortoq shear zones) which suggest that the fabrics within the nagssugtoqidian orogen may have had some influence on the fault complexity of the davis strait. the conclusions from this study show that the fault patterns and sense of movement on faults onshore reflect the stress fields that govern the opening of the labrador sea – davis strait – baffin bay seaway, and that the wrench couple on the ungava transform system played a dominant role in the development of the onshore fault patterns. acknowledgements the authors would like to thank bp (norway) and statoil (uk) for providing additional funding for this field research, and to nerc for funding r.w.w.’s ph.d. research (ner/s/s/2001/06740). thorough and insightful reviews from s. bergh and t.c.r. pulvertaft and additional help from the latter concerning numerous details are greatly appreciated. the bulletin editors, a.a. garde and f. kalsbeek, are also thanked for their helpful and encouraging comments. references bonow, j.m. 2004: palaeosurfaces and palaeovalleys on north atlantic previously glaciated passive margins – reference forms for conclusions on uplift and erosion. ph.d. thesis. thesis in geography with emphasis on physical geography 30, 17 pp. + 4 articles. stockholm university, sweden. butler, r.w.h., holdsworth, r.e. & lloyd, g.e. 1997: the role of basement reactivation in continental deformation. journal of the geological society (london) 154, 69–71. chalmers, j.a. 2000: offshore evidence for neogene uplift in central west greenland. global and planetary change 24, 311–318. chalmers, j.a. & laursen, k.h. 1995: labrador sea: the extent of continental crust and the timing of the start of sea floor spreading. marine and petroleum geology 12, 205–217. chalmers, j.a. & pulvertaft, t.c.r. 2001: development of the continental margins of the labrador sea: a review. in: wilson, r.c.l. et al. (eds): non-volcanic rifting of continental margins: a comparison of evidence from land and sea. geological society special publication (london) 187, 77–105. chalmers, j.a., dahl-jensen, t., bate, k.j. & whittaker, r.c. 1995: geology and petroleum prospectivity of the region offshore southern west greenland. rapport grønlands geologiske undersøgelse 165, 13–21. chalmers, j.a., pulvertaft, t.c.r., marcussen, c. & pedersen, a.k. 1999: new insight into the structure of the nuussuaq basin, central west greenland. marine and petroleum geology 16, 197–224. chian, d. & louden, k.e. 1994: the continent-ocean crustal transition across the southwest greenland margin. journal of geophysical research 99, 9117–9135. christiansen, f.g., bojesen-koefoed, j.a. & chalmers, j.a. 2001: petroleum geological activities in west greenland in 2000. geology of greenland survey bulletin 189, 24–33. dalhoff, f., chalmers, j.a., gregersen, u., nøhr-hansen, h., rasmussen, j.a. & sheldon, e. 2003: mapping and facies analysis of paleocene – mid-eocene seismic sequences, offshore southern west greenland. marine and petroleum geology 20, 935–986. dam, g. 2002: sedimentology of magmatically and structurally controlled outburst valleys along rifted volcanic margins: examples from the nuussuaq basin, west greenland. sedimentology 49, 505–532. dam, g. & sønderholm, m. 1998: sedimentological evolution of a fault-controlled early paleocene incised-valley system, nuussuaq basin, west greenland. in: shanley, k.w. & mccabe, p.j. (eds): relative role of eustasy, climate, and tectonism in continental rocks. society of economic paleontologists and mineralogists special publication 59, 109–121. dam, g., nøhr-hansen, h., pedersen, g.k. & sønderholm, m. 2000: sedimentary and structural evidence of a new early campanian rift phase in the nuussuaq basin, west greenland. cretaceous research 21, 127–154. de paola, n., holdsworth, r.e., mccaffrey, k.j.w. & barchi, m.r. 2005: partitioned transtension: an alternative to basin inversion models. journal of structural geology 27, 607–625. dewey, j.f. 2002: transtension in arcs and orogens. international geology review 44, 402–439. 204 dooley, t., mcclay, k. & bonora, m. 1999: 4d evolution of segmented strike-slip fault systems: applications to nw europe. in: fleet, a.j. & boldy, s.a.r. (eds): petroleum geology of northwest europe, proceedings of the 5th conference, 215–225. london: geological society. escher, j.c. & pulvertaft, t.c.r. 1995: geological map of greenland, 1:2 500 000. copenhagen: geological survey of greenland. geoffroy, j., gélard, j.p., lepvrier, c. & olivier, p. 1998: the coastal flexure of disko (west greenland), onshore expression of the ‘oblique reflectors’. journal of the geological society (london) 155, 463–473. geoffroy, j. et al. 2001: southeast baffin volcanic margin and the north american – greenland plate separation. tectonics 20, 566–584. henriksen, n., higgins, a.k., kalsbeek, f. & pulvertaft, t.c.r. 2000: greenland from archaean to quaternary. descriptive text to the geological map of greenland, 1:2 500 000. geology of greenland survey bulletin 185, 93 pp. holdsworth, r.e., butler, c.a. & roberts, a.m. 1997: the recognition of reactivation during continental deformation. journal of the geological society (london) 154, 73–78. japsen, p., bonow, j., klint, k.e.s. & jensen, f.k. 2002: neogene uplift, erosion and resedimentation in west greenland. field report summer 2002. danmarks og grønlands geologiske undersøgelse rapport 2002/71, 114 pp. japsen, p., green, p.f. & chalmers, j.a. 2005: separation of palaeogene and neogene uplift on nuussuaq, west greenland. journal of the geological society (london) 162, 299–314. larsen, j.g. & pulvertaft, t.c.r. 2000: the structure of the cretaceous–palaeogene sedimentary-volcanic area of svartenhuk halvø, central west greenland. geology of greenland survey bulletin 188, 40 pp. marker, m., mengel, f., van gool, j. and field party 1995: evolution of the palaeoproterozoic nagssugtoqidian orogen: dlc investigations in west greenland. rapport grønlands geologiske undersøgelse 165, 100–105. mcclay, k.r. 1987: the mapping of geological structures (geological society of london handbook), 161 pp. london: john wiley and sons. nieto-samaniego, a.f. & alaniz-alvarez, s.a. 1997: origin and tectonic interpretation of multiple fault patterns. tectonophysics 279, 197–206. petit, j.-p. 1987: criteria for the sense of movement on fault surfaces in brittle rocks. journal of structural geology 9, 597–608. ramberg, h. 1949: on the petrogenesis of the gneiss complexes between sukkertoppen and christianshaab, west greenland. meddelelser fra dansk geologisk forening 11, 312–327. reches, z. 1983: faulting of rocks in three-dimensional strain fields. ii. theoretical analysis. tectonophysics 95, 133–156. storey, m., duncan, r.a., pedersen, a.k., larsen, l.m. & larsen, h.c. 1998: 40ar/39ar geochronology of the west greenland tertiary volcanic province. earth and planetary science letters 160, 569–586. taylor, b., crook, k. & sinton, j. 1994: extensional transform zones and oblique spreading centres. journal of geophysical research 99(b10), 19707–19718. van gool, j.a.m., connelly, j.n., marker, m. & mengel, f.c. 2002: the nagssugtoqidian orogen of west greenland: tectonic evolution and regional correlations from a west greenland perspective. canadian journal of earth sciences 39, 665–686. whittaker, r.c. 1995: a preliminary assessment of the structure, basin development and petroleum potential offshore central west greenland. open file series grønlands geologiske undersøgelse 95/9, 33 pp., 6 maps. woodcock, n.h. & schubert, c. 1994: continental strike-slip tectonics. in: hancock, p.l. (ed.): continental deformation, 251–263. oxford: pergamon press. __________________________________________________________________________________________________________________________________________________________________________________________________________________________ manuscript received 26 october 2004; revision accepted 1 november 2005 geological survey of denmark and greenland bulletin 15, 2008, 61-64 the greenland ice sheet has been losing mass at a dramatic rate in recent years, raising political concern worldwide due to the possible impact on global sea level rise and climate dynamics (luthcke et al. 2006; rignot & kanagaratnam 2006; velicogna & wahr 2006; ipcc 2007; shepherd & wingham 2007). the arctic region as a whole is warming up much more rapidly than the globe at large (acia 2005) and it is desirable to quantify these changes in order to provide the decision-makers with a firm knowledge base. to cover this need, the danish ministry of climate and energy has now launched a new programme for monitoring of the green land ice sheet (promice), designed and operated by the geological survey of denmark and greenland (geus) in collaboration with the national space institute at the tech nical university of denmark and asiaq (greenland survey). the aim of the programme is to quantify the annual mass loss of the greenland ice sheet, track changes in the extent of local glaciers and ice caps, and track changes in the position of the ice-sheet margin. observing and modelling the ice-sheet surface-mass balance surface mass balance will be estimated using data from a new network of automatic mass-balance stations (ams) on the margin of the greenland ice sheet (fig. 1). the final station network will include a total of 14 stations located in seven climatically different regions of the greenland ice sheet and is intended to be fully operational from 2011. at each location, one station will be placed in the lower ablation zone, and another in the higher ablation zone near the equilibrium-line altitude. currently, six of these stations are operational in four regions of the ice-sheet margin (fig. 1). the promice network will complement the u.s. greenland climate network (gc-net) which consists of c. 15 stations, mainly in the accumulation zone of the ice sheet (steffen & box 2001) and the dutch k-transect of three stations on the icesheet margin near kangerlussuaq, west greenland (fig. 1; van de wal et al. 2005). this collaboration means that icea new programme for monitoring the mass loss of the greenland ice sheet andreas p. ahlstrøm and the promice project team* © geus, 2008. geological survey of denmark and greenland bulletin 15, 61–64. available at: www.geus.dk/publications/bull 61 kronprins christian land thule kangerlussuaq disko nuuk qassimiut tasiilaq scoresby sund melville bugt 250 km promice stations american stations dutch station flight path60° 70° 80° 50° 80° 70° 60° 50° 40° 30° 20° 10° 0°90° 40° 30° *peter gravesen, signe bech andersen, dirk van as, michele citterio, robert s. fausto, søren nielsen, hans f. jepsen, steen savstrup kristensen, erik lintz christensen, lars stenseng, rene forsberg, susanne hanson and dorthe petersen fig. 1. overview map showing current and future promice activities. large dots signify several stations along transects in the ablation zone of the ice sheet, small dots signify individual stations in the accumulation area. by summer 2007, six promice stations were in operation on the ice-sheet margin: two near tasiilaq, two near nuuk, one in melville bugt and one in south greenland near qassimiut. the dots near thule, kronprins christian land and inner scoresby sund signify locations of future station transects to be established over the next three years. the flight lines show the promice flight in august 2007, covering the entire margin apart from a short high elevation segment in east greenland. note that the flight paths also pass down 20 of the most significant outlet glaciers from the greenland ice sheet. sheet surface melting and its climatic causes can be effectively monitored in all geographical regions of the greenland ice sheet. the observations from the station network provide the input to perform the spatially distributed modelling of the surface melting necessary to quantify the total meltwater runoff from the greenland ice sheet. such modelling has previously relied on observational data from coastal stations in a maritime climate quite different from stations located at the ice-sheet margin and from stations in the accumulation zone, complemented only by the measurements from the few u.s. and dutch stations on the ice-sheet margin in west green land (box et al. 2006). the innovative geus station setup includes measurement of air temperature, relative humidity, wind speed and direction, atmospheric pressure, incoming and outgoing shortand long-wave radiation, ice temperature, surface velocity, snow depth and ice ablation (fig. 2). the climate and ice temperature data allow calculation of the complete surface-energy balance. together with the snow depth measurement, these data make it possible to determine the specific cause of an observed change in surface melt. this implies that we will be able to answer questions such as whether the impact on ice-sheet ablation of rising air temperatures is offset by a corresponding increase in snow precipitation. the data are stored locally in the data logger, but are also transmitted by satellite to geus at regular intervals during the melt season, when solar panels deliver the necessary additional power and the data are most crucial. figure 3 shows some samples of transmitted data. station power and tilt are also monitored in order to minimise costly maintenance visits. generally, station visits by helicopter constitute the main expense in operating the network on the ice-sheet margin. this fact has prompted the development of a new type of ablation sensor at geus, based on measuring the pressure of anti-freeze liquid at the end of a flexible, closed hose in serted into the ice to a depth of 30–40 m. as the ice melts, the length of the hose in the ice diminishes, lowering the pressure exerted by the liquid column on the sensor at the bottom of the hose. the top of the hose is connected to a liquid reservoir at a fixed level on the station, making it possible to deduce ice-sheet surface melt from pressure change. an early version of the system has been described by bøggild et al. (2005). once this system is operational, it should in theory be possible to leave the station unattended for as long a time as it takes the hose to melt out of the ice, i.e. more than 5 years, as the remaining part of the station is ‘floating’ on the ever-changing ice surface. in practice, stations can rarely be left unattended in the harsh environment on the ice-sheet margin for that long, and most instruments would in any case require calibration or maintenance at more regular intervals. however, even a reduction of visits from every year to every second year is substantial. currently, the pressure-sensor system is supplemented by a more traditional set up with a sonic ranging device mounted on stakes drilled into the ice. the 62 1 2 3 4 9 8 7 6 5 8 10 12 13 14 11 fig. 2. the new promice automatic mass-balance station, in this case the station with the lowest elevation on the greenland ice-sheet margin near tasiilaq, east greenland. 1, short-wave radiation (in/out); 2, longwave radiation (in/out); 3, station tilt; 4, satellite transmitter; 5, wind speed and direction; 6, snow height (sonar); 7, air temperature and relative humidity; 8, ablation (pressure sensor); 9, solar panel; 10, dogger logger enclosure; 11, gps and multiplexer enclosure; 12, battery box; 13, ice temperature 0–10 m (thermistors); 14, ablation (sonar and stakes). 500 a b c 0.2 0.0 –0.2 –0.4 300 100 –100 38.505 38.510 38.500 la tit ud e (d eg re e) 45 0 m su rf ac e he ig ht ch an ge ( m ) n et r ad ia tiv e flu x (w m –2 ) 38.495 1 september 15 september fig. 3. a: net radiative flux measured by a cnr-1 instrument, used on the promice stations in 2007. the cnr-1 measures incoming and outgoing shortand long-wave radiation and enables calculation of the surface albedo. b: snow depth and ablation, measured respectively by a sr-50 sonic ranging device installed on the station tripod and on stakes drilled into the ice. c: latitudinal movement of the station recorded by the gps. all graphs have been produced from satellite-transmitted data received at geus. sonic-ranger system requires annual visits for re-drilling of stakes, but is necessary as long as the pressure-sensor system is not completely reliable. both systems record ice melt on an hourly basis, yielding the perfect validation data for determining the individual energy-balance components from the climate data collected. quantifying the mass loss caused by iceberg calving iceberg calving from the outlet glaciers of the greenland ice sheet, often termed the ice-dynamic mass loss, is responsible for most of the acceleration in the mass loss during the last decade (rignot & kanagaratnam 2006). to quantify this part of the mass loss, we combine an airborne survey yielding surface elevation and ice depth across the entire ice-sheet margin, with ice-sheet surface velocity derived from satellite radar. the ice-dynamic mass loss is then derived by calculating the ice flux from the interior of the ice sheet towards the outlet glaciers, while correcting for the surface melt between the flux gate and the calving front of the glacier (rignot & kanagaratnam 2006). the route chosen complements the airborne ice-sheet elevation measurements carried out by u.s. researchers. the route was designed to cover the accumulation zone and the centre line of the main outlet glaciers of the greenland ice sheet (krabill et al. 2004). the first airborne survey within promice was carried out in august 2007, measuring ice-sheet elevation with a riegl scanning laser altimeter (forsberg et al. 2001) and ice-sheet thickness with a 60 mhz coherent ice-penetrating radar (christensen et al. 2000). an example of the radar data is shown in fig. 4. the airborne survey will be repeated every two or three years to monitor temporal changes in the elevation of the greenland ice-sheet margin. the radar had difficulties penetrating the ice sheet in the southernmost part of greenland, probably due to the extensive surface melting experienced in 2007 and the fact that the survey was carried out in august at the peak of the melt season. the time of year had been selected to minimise the influence of snow depth on the elevation measurements, but may have to be shifted to the spring for future surveys in order to obtain the ice-sheet thickness. a new p-band radar is currently being developed at the technical university of denmark for the european space agency. this may replace the older 60 mhz radar on future surveys, if it is more successful in retrieving bedrock returns beneath the ice sheet and is made available for our use. monitoring the change of glaciers and ice caps in greenland in addition to the antarctic and greenland ice sheets, melting of glaciers and ice caps around the world is contributing significantly to the present-day sea-level rise, and yet almost nothing is known about the current state of the c. 20 000 glaciers and ice caps in greenland. to remedy this lack of knowledge, promice aims at delivering glacier and icesheet outlines to the global land ice measurements from space (glims) project, in order to compile a complete in ventory of greenland glaciers and ice caps as well as the icesheet margin. representative glaciers can be selected from the inventory for detailed volume-change studies of the reaction to climate change and used for up-scaling to larger regions. the island of disko in west greenland was chosen as a test region to develop the methodology for a number of reasons: (a) vectorised maps and elevation contours are available from an airborne survey in 1985; (b) the island is covered by a recent landsat 7 etm+ satellite image, facilitating a comparison over a 16-year period; (c) the island is covered by the glacier inventory of west greenland (wggi) (weidick et al. 1992); and (d) the island exhibits a wide range of glacier types and problems representative for greenland as a whole. the wggi covers more than 5000 glaciers, and is published as a map compilation including a table summarising glacier specific data, such as area, elevation span and orientation. a digital version of the table was checked for errors and included in the geographic information system, arcgis. glacier positions needed manual correction due to the limited precision listed in the inventory table. the relevant inventory maps were scanned, geo-referenced and likewise imported into arcgis, in order to yield additional information on glacier extent during the little ice age (lia) as derived from aerial photographs. a landsat 7 etm+ image from 2001 available from the global land cover facility was utilised to obtain a more recent inventory of the glaciers on 63 2 km 200 m surface of ice surface of bedrock fig. 4. data from an ice-penetrating radar. the data have been postprocessed and interpreted to determine the base and surface of the ice sheet at a site in west greenland. eventually the ice thickness is determined using the radar reflection from the bedrock in combination with laser altimeter measurements of the ice-sheet surface. 64 disko, using the glims algorithm (paul & kääb 2005). a 50 m resolution digital elevation model was created from the 100 m elevation contours of the 1985 map, to assist in determining ice-flow divides and orientation controls on glacier sensitivity to climate change. preliminary results show that in the period 1985–2001, glaciers on disko retreated by 8%, whereas no significant areal change could be determined (<0.5%). it was also found that the wggi could not be used for determination of glacier area change since the lia, but that the glaciers on the island of disko had retreated 44% between the lia and 2001. the initial results thus point to at a more limited reaction to recent warming trends in west greenland, compared to the strongly responding glaciers and ice caps in alaska (arendt et al. 2002). outlook the data acquired by the monitoring programme will be stored in a database constructed for the purpose at geus. so far, the process of defining the database requirements and user needs has been initiated. once the monitoring programme is operational, data will be made available within a year from acquisition through this database. already at this early stage, promice has been woven into a number of research activities, national as well as international, utilising the existing field operations as a vehicle for scientific projects. through the extensive monitoring programme, geus has become a natural partner in most international glaciological research activities on the greenland ice sheet. the effort is also a strong asset for current and future hydropower investigations by the greenland home rule, as well as for prospecting activities near the ice margin. over the next three years, the full station network will be established, the ice-sheet surface velocity software will be fi nal ised and the glaciological modelling tools for estimating mel t ing and dynamic mass loss will be developed and applied. with the new monitoring programme of the greenland ice sheet, we will not only be able to provide an answer to how much the green land ice sheet contributes to global sea-level rise, but also make it possible for the global research community to gain access to key data sets to conduct their own investigations. acknowledgement the programme for monitoring of the greenland ice sheet (promice) is partly funded by dancea under the danish ministry of climate and energy. references acia 2005: arctic climate impact assessment, 1042 pp. cambridge: cambridge university press. arendt, a.a., echelmeyer, k.a., harrison, w.d., lingle, c.s. & valentine, v.b. 2002: rapid wastage of alaska glaciers and their contribution to rising sea level. science 297, 382–386. bøggild, c.e., olesen, o.b., ahlstrøm, a.p. & jørgensen, p. 2005: auto matic glacier mass balance observations using pressure sensors. jour nal of glaciology 50(169), 303–305. box, j.e., bromwich, d.h., veenhuis, b.a., bai, l.-s., stroeve, j.c., rogers, j.c., steffen, k., haran, t. & wang, s.-h. 2006: greenland ice sheet surface mass balance variability (1988–2004) from calibrated polar mm5 output. journal of climate 19, 2783–2800. christensen, e.l., reeh, n., forsberg, r., jørgensen, j.h., skou, n. & woelders, k. 2000: a low cost glacier mapping system. journal of glaciology 46(154), 531–537. forsberg, r., keller, k. & jacobsen, s.m. 2001: laser monitoring of ice elevations and sea-ice thickness in greenland. international archives of photogrammetry and remote sensing 34, 163–169. ipcc 2007: climate change 2007: the physical science basis. contribution of working group i to the fourth assessment report of the inter governmental panel on climate change [solomon, s. et al. (eds)]. cam bridge: cambridge university press. krabill, w. et al. 2004: greenland ice sheet: increased coastal thinning. geophysical research letters 31, l24402, doi:10.1029/2004gl021533. luthcke, s.b., zwally, h.j., abdalati, w., rowlands, d.d., ray, r.d., nerem, r.s., lemoine, f.g., mccarthy, j.j. & chinn, d.s. 2006: recent greenland ice mass loss by drainage system from satellite gravity observations. science 314, 1286–1289. paul, f. & kääb, a. 2005: perspectives on the production of a glacier inventory from multispectral satellite data in the canadian arctic: cumberland peninsula, baffin island. annals of glaciology 42, 59–66. rignot, e. & kanagaratnam, p. 2006: changes in the velocity structure of the greenland ice sheet. science 311, 986–990. shepherd, a. & wingham, d. 2007: recent sea-level contributions of the antarctic and greenland ice sheets. science 315, 1529–1532. steffen, k. & box, j.e. 2001: surface climatology of the greenland ice sheet: greenland climate network 1995 –1999. journal of geophysical research, 106(d24), 33,951–33,964. van de wal, r.s.w., greuell, w., van den broeke, m.r., reijmer, c.h. & oerlemans, j. 2005: surface mass-balance observations and automatic weather station data along a transect near kangerlussuaq, west green land. annals of glaciology 42, 311–316. velicogna, i. & wahr, j. 2006: acceleration of greenland ice mass loss in spring 2004. nature 443, 329–331. weidick, a., bøggild, c.e. & knudsen, n.t. 1992: glacier inventory and atlas of west greenland. rapport grønlands geologiske undersøgelse 158, 194 pp. authors’ addresses a.p.a., p.g., s.b.a., d.v.a., m.c., r.s.f., s.n. & h.f.j., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: apa@geus.dk s.s.k., e.l.c., l.s., r.f. & s.h., national space institute, technical university of denmark, ørsteds plads, bldg. 348, dk-2800 kgs. lyngby, denmark. d.p., asiaq – greenland survey, p.o. box 1003, qatserisut 8, dk-3900 nuuk, greenland. geological survey of denmark and greenland bulletin 17, 2009, 29-32 the møns klint glaciotectonic complex (fig. 1) exposed in the n–s-trending chalk cliff on the east coast of the island of møn in south-east denmark is one of the most famous glaciotectonic geosites in the world. people of all nationalities are attracted to the site, which has more than 300 000 visitors per year. many of them may not realise the uniqueness of the glaciotectonic framework, and are probably more fascinated by the spectacular view of the white cliff and chalk peaks separated by the deep green gorges. however, without the glaciotectonic deformation the cliffs would never have formed. instead the cretaceous chalk would still have been resting horizontally below the seabed, covered by glaciofluvial sand, gla ciolacustrine clay and clayey till. in the summer 2007, a new natural science exhibition centre, geocenter møns klint, was opened. the exhibition fo cuses on the geology of denmark with special reference to the chalk cliffs of møns klint. prior to the decision to build the centre, the geological survey of denmark and greenland was asked to provide an evaluation of the landslide risk for the site, because landslides regularly occur along the cliff section (pe dersen 2003). a detailed structural analysis was consequently carried out at the planned site of the centre just above the maglevandsfald gorge (fig. 1). during this investigation it became obvious that understanding the glaciotectonic framework was a prerequisite for the geological risk analysis. thus the structural details at maglevandsfald became a key point for the glaciotectonic model of møns klint which we present in this paper. geological setting of møns klint the cretaceous chalk represents the oldest bedrock affected by glaciotectonic thrusting in denmark. the chalk consists of a fine-grained matrix of coccoliths and remains from invertebrates (surlyk & håkansson 1999). the chalk at møn is correlated with the upper part of the tor formation in the north sea and the danish basin (surlyk et al. 2003). an erosional unconformity on top of the chalk represents a hiatus of about 60 million years. due to erosion, the unconformity is found at gradually lower and lower levels towards the southern part of the cliff, which indicates that the pre-qua ternary relief probably contributed to the formation of glaciotectonics in the area. in some parts of møn the pre-qua ternary unconformity is overlain by a saalian till, marine eemian clay and three weichselian tills intercalated by glaciolacustrine and glaciofluvial deposits (houmark-nielsen 1994). however, at møns klint an old weichselian till is the oldest quaternary deposit recognised at present (pedersen & gravesen 2006). this till is a grey and reddish sandy till that grades up into a gravelly, stone-rich top surface indicating terrestrial conditions. we tenta tively correlate this unit with the ristinge klint till (houmarknielsen 1987, 1994), which was deposited by a baltic ice stream in the early part of the weichselian (c. 60 000 years bp) from a source area in the baltic sea. when the baltic ice stream melted back the depression became occupied by a huge lake, in which dark glaciolacustrine mud rich in dropstones was deposited (houmark-nielsen 1994; houmark-nielsen & kjær 2003). structural development of maglevandsfald: a key to understanding the glaciotectonic architecture of møns klint, se denmark stig a. schack pedersen and peter gravesen © geus, 2009. geological survey of denmark and greenland bulletin 17, 29–32. available at: www.geus.dk/publications/bull 29 fig. 1. aerial photograph of møns klint. the highest cliff in the centre is dronningestolen. to the left the maglevandsfald gorge leads from the location of geocenter møns klint down to the beach. at the left side of dron ningestolen a wedge-shaped structure is seen (red arrow), which was formed in the first deformation phase of the glaciotectonic complex. the thrust structures to the right of dronningestolen were thrust to the sw during the last deformation phase. the location of møns klint is shown on the inset map of denmark. dronningestolenmaglevandsfald geocentre rosa_2008:rosa-2008 01/07/09 15:48 side 29 the unit is informally called the elephant clay due to its characteristic erosion features. at the late weichselian glacial maximum the swedish ice advanced over the eastern part of the danish basin onto the main stationary line depositing the mid danish till on møn (fig. 2; houmark-nielsen 1987; pedersen & gravesen 2006). at about 18 000 years bp the young baltic ice stream reached møn (houmark-nielsen & kjær 2003). the pressure from the ice was directed away from the central axis of the ice stream towards both north and south, where it thrust up the chalk sheets at møns klint. the glaciotectonic structures at møns klint can be divided into three architectural elements. in the southern part the thrust sheets form an imbricate fan with thrust faults striking e–w, in the central part the structures form an antiformal stack, and farthest to the north the thrust sheets dip gently to the south in the foreland regime of the thrust deformation (fig. 3; pedersen 2000). towards the end of the weich selian the scandinavian ice sheet melted back. however, a re advance from southern sweden reached møn from eastnorth-east at a time between 17 000 and 15 000 years bp, which was responsible for superimposed deformation of the glaciotectonic complex (pedersen 2000). photogrammetric mapping and a borehole investigation a 3-d terrain model of the area was developed for the risk evaluation. the data for the model were provided by a photo grammetric survey and a digital model was constructed, in which the trace of the glacial deposits, interbedded between the unconformity on top of the chalk and the base of an overthrust chalk sheet, was positioned (fig. 4). in addition, a borehole was drilled to investigate whether karst cavities or an unforeseen thrust-fault zone with dip towards the cliff and beach were present. if a fault zone exists below the location it could be a candidate for a landslip similar to the one that destroyed the store taler chalk peak in january 2007 (fig. 5). the borehole was placed on top of the cliff 100 m above sea level close to the old hotel store klint (fig. 6). the uppermost 16.3 m consist of glacial sediments resting upon the chalk unconformally. the oldest unit in the glacial succession is a 1 m thick till that is correlated with the ristinge klint till. this formation is overlain by 2.3 m glaciolacustrine mud, which grades into the nearly 2 m thick mid danish till. a 7.7 m thick unit of glaciofluvial sand and gravel overlies the mid danish till; the upper part of the glacial succession is a 3.5 m thick, sandy till that was deposited by the young baltic ice. below the unconformity the drilling penetrated 48 m of cretaceous chalk, after which the drilling was 30 s n 100 m sea level glaciofluvial sand mid danish till glaciolacustrine clay ristinge klint till maastrichtian chalk thrust fault 10 0 m fig. 2. stratigraphical and structural framework of the southern imbricate fan at møns klint. the thickness of the lithological units is variable, and especially the glaciofluvial sand above the mid danish till increases in thickness, where the sand was deposited in piggyback basins during the thrust-fault deformation. from pedersen & gravesen (2006). imbricate fan antiformal stack ramp slotsgavle foreland thrust dronningestolen antiformal stack maglevandsfald ramp gråryg imbricate ene jydelejet superimposed thrusting sse young baltic ice advance fig. 3. principal structural framework of the møns klint glaciotectonic complex. blue: chalk; brown and orange: quaternary deposits. red lines outline the thrust faults. to the south the e–w-trending ridges represent an imbricate fan, and in the central part an antiformal stack is responsible for the formation of dronningestolen. two principal sketches of the structures are shown in the top left corner. in the distal part of the complex to the north a gently dipping imbricate fan was formed in the nearforeland regime. the structures from the distal regime to maglevandsfald were strongly affected by superimposed deformation caused by a re advance of ice from skåne. after pedersen (2000). rosa_2008:rosa-2008 01/07/09 15:48 side 30 unfortunately stopped for logistic reasons. it would have been perfect if the borehole had reached the level of the lower thrust fault and the underlying glacial succession. however, the data demonstrate that neither a karst cavity nor an eastward dipping thrust-fault zone was present. thrust-fault tectonics and superimposed deformation the structures in the møns klint glaciotectonic complex are directly comparable to structures in thin-skinned thrustfault belts formed by gravity spreading (pedersen 2005). the geomorphological expression of a mountain range is clearly seen in the parallel-ridged landscape of høje møn (the eastern, hilly part of møn). for detailed structural analysis five cross-sections were constructed across the maglevandsfald (figs 6, 7). the data for the cross-sections were taken from the photogrammetric survey combined with a structural investigation of the cliff section immediately south of maglevandsfald. in the cross-sections two horizons of glacial successions are shown (fig. 7). the lower glacial horizon rests on the unconformity on top of the chalk. the top of the glacial deposits is bordered by a thrust fault, where the base of the chalk sheet can be classified as a hanging-wall flat related to the earliest glaciotectonic deformation. farther to the south this flat can be followed into a hanging-wall ramp, which dips into the subsurface below the beach. this ramp strikes e–w, indicating that it is part of the antiformal stack structure (fig. 3). as the décollement zone must be situated at the base of the chalk unit, which in fig. 7 can be seen to be about 60 m thick, the depth to this zone must be of similar magnitude, probably about 75 m below sea level. therefore the chalk below the lower glacial unit is also interpreted as a thrust sheet lifted up from the décollement zone. the lower glacial unit is folded in a sw-verging overturned structure, which has been deformed by the second glaciotectonic event. the structure is partly hidden by soil and vegetation in the maglevandsfald gorge, but the upper limb of the structure is exposed at the cliff edge of magle vandsnakken (fig. 7), where the unit continues into the wedge-shaped feature in the southern part of the dron ningestolen cliff section (fig. 1). the wedge-shaped feature was formed during the foreland-dipping thrusting of the antiformal stack (compare with fig. 3). the thrust-fault flat 31 fig. 4. 3-d model of the maglevandsfald gorge. the position of the ex ploratory borehole dgu 228.84 is indicated with a dot, located be tween the exhibition centre and the old hotel. the upper and lower bound aries of the glacial succession are shown with a brown line. contour interval 10 m. steps to the beach 100 m a.s.l. glacial succession dgu 228.84 n fig. 5. the landslide at store taler took place at the end of january 2007. a volume of more than 100 000 m3 was displaced along an inherited thrust-fault surface formed during the last phase of deformation at møns klint. store taler is located in the northern part of møns klint. location of cross-sections dronningestolen dgu 228.84dgu 228.84 parking place parking place 1 2 3 4 5 hotel geocentre hotel geocentre 100 m thrust unconformity n 10 20 30 50 60 4050 60 708090 95 100 130 120 110 100 90 80 70 10 20 30 50 60 4050 60 708090 95 100 130 120 110 100 90 80 70 fig. 6. the maglevandsfald area with location of the exploratory borehole mentioned in the text and the cross-sections applied in the structural analysis. rosa_2008:rosa-2008 01/07/09 15:48 side 31 32 above the glacial unit is folded in the maglevandsfald gorge structure with a fold axis trending nw–se, which demonstrates the superimposed deformation. a similar wedge-shaped structure is present just below the edge of the dronningestolen cliff, which may indicate that a thrust sheet existed even higher in the antiformal stack before the young baltic ice truncated the structure during the advance towards nnw. final remarks møns klint is one of the most famous glaciotectonic localities in the world. the structures in the møns klint glacio tec tonic complex are comparable to structures in thinskinned thrust-fault belts formed by gravity spreading. the glaciotectonic complex is responsible for the parallel-ridge landscape known as høje møn, where aborre bjerg (143 m) represents the highest hill in eastern denmark. a combination of an imbricate fan and an antiformal stack is respon sible for the impressive tectonic complex. its main architecture was formed during the young baltic ice stream advance dated to about 18 000 years bp, but a readvance of ice from southern sweden from 17 000 to 15 000 years bp is responsible for the superimposed deformation of the complex. the most impressive unit forming the thrust sheets is the maastrichtian chalk. the chalk is part of the huge carbonate platform that spread over northern europe during the late cretaceous. the top surface of the chalk, the pre-quaternary unconformity, was originally situated more than 20 m below sea level, and the décollement zone for the thrusting is probably located about 80–100 m below sea level in a marl-rich variety of the chalk. references houmark-nielsen, m. 1987: pleistocene stratigraphy and glacial history of the central part of denmark. bulletin of the geological society of denmark 36, 1–189. houmark-nielsen, m. 1994: late pleistocene stratigraphy, glacial chronology and middle weichselian environmental history from klintholm, møn, denmark. bulletin of the geological society of denmark 41, 181–202. houmark-nielsen, m. & kjær, k. 2003: southwest scandinavia, 40–15 kyr bp: palaeography and environmental change. journal of quaternary science 18, 769–786. pedersen, s.a.s. 2000: superimposed deformation in glaciotectonics. bulletin of the geological society of denmark 46, 125–144. pedersen, s.a.s. 2003: vurdering af skredrisiko for området oven for magle vandsfaldet på møns klint. strukturel undersøgelse af de glacialtektoniske forhold i klintområdet ved hotel storeklint, møns klint, møn. danmarks og grønlands geologiske undersøgelse rapport 2003/50, 31 pp. pedersen, s.a.s. 2005: structural analysis of the rubjerg knude glaciotectonic complex, vendsyssel, northern denmark. geological survey of denmark and greenland bulletin 8, 192 pp. pedersen, s.a.s. & gravesen, p. 2006: geological map of denmark, 1:50 000, møn. copenhagen: geological survey of denmark and greenland. surlyk, f. & håkansson, e. 1999: maastrichtian and danian strata in the southeastern part of the danish basin. in: pedersen, g.k. & clemmensen, l.b. (eds): ias field trip guidebook. contribution to geology, 29–58. copen hagen: geological museum, university of copenhagen. surlyk, f., dons, t., clausen, c.k. & higham, j. 2003: upper cretaceous. in: evans, d. et al. (eds): the millennium atlas: petroleum geology of the central and northern north sea, 213–233. london: the geological society of london. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: sasp@geus.dk 140 120 100 80 60 40 20 0 140 120 100 80 60 40 20 h ei gh t ( m a .s. l.) h ei gh t ( m a .s. l.) h ei gh t ( m a .s. l.) 140 120 100 80 60 40 20 0 maglevandsfald maglevandsfald cross-section 1 cross-section 3 maglevandsfald maglevandsnakken maglevandsnakken glacial succession on top of upper thrust sheet thrust fault glacial succession at the base of the cliff cross-section 5 dronningestolen nesw 1000 200 300 400 1000 200 300 400 1000 200 300 400 distance (m) borehole 228.84 fig. 7. three of the cross-sections across the maglevandsfald gorge (for location see fig. 6). the cross-sections were constructed perpendicular to the fold axis of the last deformation, which created a recumbent anticline and syncline pair below a thrust fault with a displacement of 20 m. the folding is outlined in the glacial units below the thrust-fault flat of the first deformation, which is responsible for the wedge-shaped feature in the top of the fold. the strike of the first deformation thrust is oblique to the second phase fold axis, and therefore the position of the wedge changes from cross-section to cross-section. in cross-section 5 the tip of the thrust wedge touches the cliff edge, which corresponds to the exposure of the wedge seen in fig. 1 immediately to the right of magle vandsfald. rosa_2008:rosa-2008 01/07/09 15:48 side 32 geological survey of denmark and greenland bulletin 20, 2010, 55–58 55 the dynamics of the large outlet glaciers in greenland is attracting both scientific and political attention due to the possible implications of a rising global sea level. extensive glaciological and meteorological monitoring programmes have been implemented to quantify and track changes in the ice sheet and local glaciers (ahlstrøm et al. 2008). the dynamic processes controlling the flow of the outlet glaciers are complex and poorly understood, involving a wealth of parameters such as bed conditions, hydrology and meteorological conditions. it is desirable to obtain as many fundamentally independent data sets as possible to understand and eventually predict the behaviour of the outlet glaciers. some processes related to ice dynamics can be detected seismologically and thus completely independently from classical ice-monitoring techniques, such as satellite remote sensing, global positioning system (gps) geodesy and automatic weather stations. detectable cryo-seismological events include high-frequency ice quakes (anandakrishnan & bentley 1993; harrison et al. 1993), calving events (o’neel et al. 2006; nettles et al. 2008) and less well understood processes such as low-frequency glacial earthquakes (ekstrom et al. 2003; nettles et al. 2008)) and glacial rumblings (rial et al. 2009). changes in ice load along the margin of the ice sheet can lead to earthquakes from glacial rebound, and earthquakes can provide an independent constraint on ice mass redistribution (johnston 1987; stewart et al. 2000; lund & näslund 2009). the greenland ice sheet monitoring network (glisn) project will monitor changes in glacier dynamics using a large broadband seismological network. the network will also improve the detection of tectonic earthquakes in greenland, thereby establishing a better baseline for local seismicity. the baseline will allow detection of future changes in seismicity caused by changes in ice load. it is the objective of the project to contribute significantly to understanding the dynamics of the greenland ice sheet and glaciers by studying cryo-seismological processes. installing and operating a large real-time seismological network in greenland is logistically complicated and expensive. an international team consisting of researchers from 10 institutions in 8 countries in europe, north america and asia are working together to meet this challenge (fig. 1). glacial earthquakes and rumblings cryo-seismological events such as glacial earthquakes and rumblings can be linked to large-scale glacier dynamics. glacial earthquakes are produced at large outlet glaciers and appear to be associated with large calving events (amundson et al. 2008; nettles et al. 2008). however, the processes leading greenland ice sheet monitoring network (glisn): a seismological approach trine dahl-jensen, tine b. larsen, peter h. voss and the glisn group fig. 1. map of broadband seismographs in greenland. 1–4: permanent stations operated by geus in cooperation with other institutions. 5–7: long-term stations maintained by geus. 8–9: new glisn long-term stations run by eth (switzerland). 10: a long-term station run by geofon (germany). 11: a permanent station at alert in canada run by iris (usa). 12: a temporary station run by geus. 13–18: planned and funded new glisn stations to be installed by iris and eth. planned station 1 9 18 8 7 12 11 5 2 17 10 15 14 13 6 4 16 3 exisiting station glacial earthquake jakobshavn isbræ kangerlussuaq gletscher 400 km 80° 70° 50° w 40° w © geus, 2010. geological survey of denmark and greenland bulletin 20, 55–58. open access: www.geus.dk/publications/bull 5656 to a glacial earthquake are still poorly understood, as are the changes in dynamics following an earthquake. glacial earthquakes are slow, low-frequency events that can be detected on seismographs worldwide. the vast majority of glacial earthquakes in greenland occur at the large outlet glaciers in western and eastern greenland (fig. 1; ekstrom et al. 2003). during a period with warmer than average temperatures in greenland, a sharp increase in the number of glacial earthquakes has been observed (ekstrom et al. 2006); during the same period there was no change in the number of observed tectonic earthquakes. the glacial earthquakes can be registered at teleseismic distances, and the teleseismic signals are sufficiently strong to locate the earthquakes. however, the signals from glacial earthquakes are fundamentally different from the signals generated by tectonic earthquakes. the short-period signal from a 4.9 magnitude glacial earthquake is barely discernable on seismograms c. 1000 km away as shown in fig. 2a. the data record has been filtered from 0.5 to 1.3 sec where p (primary) waves normally dominate the seismogram. close to the epicentre of a glacial earthquake, the p wave is clearly visible (at sødalen and summit; fig. 2a), whereas it requires supporting stations to identify the signal at danmarkshavn approximately 974 km away and the short-period signal is completely lost at alert 1645 km away. for a tectonic earthquake of similar magnitude, the short-period p wave can travel many thousands of kilometres without being dissipated below the noise level. the longer periods in the signal from a glacial earthquake are not dissipated as rapidly as the shorter periods (fig. 2b) and retain a good quality at teleseismic distances. the higher-frequency waves generated by a glacial earthquake contain information about processes in and around the glacier during the earthquake. in order to understand the earthquake processes it is necessary to model waveforms recorded at local and regional distances where the full frequency range is retained. this is currently only possible for a small portion of the glacial earthquakes occurring in greenp p 20 40 60 20 40 alert (station 11) danmarkshavn (station 2) summit (station 10) sødalen (station 16) a date: 20 october 2000 20 40 60 20 40 alert (station 11) danmarkshavn (station 2) summit (station 10) sødalen (station 16) 1645 km 974 km 442 km 113 km b fig. 2. a large glacial earthquake released by the kangerlussuaq gletscher (fig. 1). the window shows 2 hours of data from four broadband seismographs in greenland and canada, located 113–1645 km from the earthquake. station numbers refer to fig. 1. the curves show vertical components band-pass filtered from 0.769–2.000 hz (a) and 0.01–0.03 hz (b). the glacial earthquake was registered from c. 02:00–02:15 utc and had a magnitude of 4.9 on the richter scale according to ekstrom et al. 2003. in a the arrival of the p wave is marked by arrows on the two closest stations (summit and sødalen). the arrival of the p wave is commonly only visible at higher frequencies and at stations close by. the data from sødalen are from a station that was in operation from 2000 to 2002. 57 land, but with the implementation of the glisn network the glacial earthquakes can be investigated in greater detail than today. glacial rumblings are slow, high-frequency events at the large outlet glaciers, and also appear to be related to calving that can be detected only at local to regional distances (fig. 3). rumblings have so far been described in only one publication for the jakobshavn isbræ (rial et al. 2009), but can be found in data from other regions of greenland as well. it has been postulated that up to 30% of the annual iceberg discharge from jakobshavn isbræ can be related to rumblings that are registered on seismographs (rial et al. 2009), but it has not yet been possible to make a similar estimate for the glacial earthquakes. the glisn project will establish and run a dense network that can capture events both at currently seismologically active glaciers and at glaciers farther to the north should the activity migrate northwards. the seismological signals may provide an early warning of changes in glacier activity. establishing the network the aim of the glisn project is to cover all of greenland and surrounding areas with as regular a seismological network as feasible (fig. 1). the stations will be upgraded to a common standard and will provide real-time data online. stations situated in greenland communities will transmit data by adsl broadband, while data from the remote stations will be transmitted by satellite, using the iridium system. new sites are chosen to cover areas close to sites where glacial seismological events are known to occur, and to ensure monitoring of the whole of greenland, so that changes in occurrence patterns such as a northwards shift can be detected. the main challenge is the operation of stations in remote areas, not least those on the greenland ice sheet. two sites on the ice sheet (sites 12 and 14; fig. 1) will be equipped with surface seismographs dug into the snow. since these are quite sensitive to settling snow and loss of levelling, these sites will also be equipped with seismographs in 300 m deep boreholes, in which the seismograph is fixed to ice instead of snow making it more stable. the network upgrade and installation were initiated in 2009, and are planned to be completed in 2011. international cooperation glisn is an international cooperation between incorporated research institutions for seismology (iris) in the usa, geological survey of denmark and greenland (geus), geoforschungszentrum network (geofon) in germany, eth zürich in switzerland, istituto nazionale di geofisica e vulcanologia (ingv) in italy, national institute of polar research (nipr) and japan agency for marine-earth science and technology (jamstec) both in japan, norsar in norway and geological survey of canada (gsc). the group is open and other participants are welcome. the operation of broadband stations in greenland builds on work by geus (dahl-jensen et al. 2003). the contribution from geus to establish glisn is the existing network, consisting of four permanent and four temporary stations, which has been in operation for up to a decade. since 2002, our german colleagues have operated the only long-term station on the ice sheet at summit camp (station 10, fig. 1), which is now also part of glisn. our swiss colleagues installed two new glisn stations in north-western greenland in 2009 and will add one more in 2010. furthermore, our colleagues from the usa have obtained a three-year grant from the national science foundation, providing the bulk of the funding for the project. this grant allows installation of five new stations and upgrading of the existing stations to 35 40 45 50 time (min) north component east component vertical component date: 14 may 2006 fig. 3. rumbling from jakobshavn isbræ observed on the broadband seismograph at station 1 (fig. 1). the window shows 20 minutes of seismic data that have been band-pass filtered from 5 to 9.9 hz. the glacier rumbling was observed from c. 12:41 to 12:44 utc and is described by rial et al. 2009. on our local scale the event was of magnitude 2.1. 5858 meet the technical standards for equipment and installation agreed for glisn – including provision of real-time data. expected outcome with glisn we seek to increase our understanding of the processes governing the ice sheet and the dynamics of the outlet glaciers. the complex processes and dynamics must be studied from different angles, such as satellite data acquisition and geodetic and glaciological observations. seismology, and thus the ability to detect events deep within and below the ice, is a new player in this field, raising many new questions. local seismicity in greenland has not been addressed in detail for many years (gregersen 1989; dahl-jensen 1984). however, increasingly detailed data have been gathered since 2000 (dahl-jensen et al. 2003) and data acquisition is rapidly increasing after the inauguration of glisn. the pressure from ice sheets usually dampens earthquakes beneath the ice (johnston 1987; stewart et al. 2000; lund & näslund 2009). if the ice sheet thins and retreats in the future, the frequency of earthquakes related to glacial rebound will probably increase. seismology is thus a tool to detect changes in the pressure on the subsurface due to climate changes. glisn also allows the continuation and expansion of the use of tectonic earthquakes to investigate the structure of the greenland crust, lithosphere and upper mantle in increasing detail (dahl-jensen et al. 2003; darbyshire et al. 2004; kumar et al. 2005, 2007; larsen et al. 2006; ucisik et al. 2008). acknowledgements the entire glisn group forms the backbone of glisn. the website www.glisn.info provides additional information on contributions and funding. references ahlstrøm, a.p. & the promice project team 2008: a new programme for monitoring the mass loss of the greenland ice sheet. geological survey of denmark and greenland bulletin 15, 61–64. amundson, j.m., truffer, m., lüthi, m.p., fahnestock, m., west, m. & motyka, r.j. 2008: glacier, fjord, and seismic response to recent large calving events, jakobshavn isbræ, greenland. geophysical research letters 35, l22501, doi:10.1029/2008gl035281. anandakrishnan, s. & bentley, c.r. 1993: micro-earthquakes beneath ice streams b and c, west antarctica: observations and implications. journal of glaciology 39, 455–462. dahl-jensen, t. 1984: jordskælv og skorpestruktur i grønland, 73 pp. unpublished m.sc. thesis, københavns universitet, danmark. dahl-jensen, t., larsen, t., wölbern, i., bach, t., hanka, w., kind, r., gregersen, s., mosegaard, k., voss, p. & gudmundsson, o. 2003: depth to moho in greenland: receiver function analysis suggests two proterozoic blocks in greenland. earth and planetary science letters 205, 379–393. darbyshire, f.a., larsen, t.b., mosegaard, k., dahl-jensen, t., gudmundsson, o., bach, t., gregersen, s., pedersen, h. & hanka, w. 2004: a first detailed look at the greenland lithosphere and upper mantle using rayleigh wave tomography. geophysical journal international 158, 267–286. ekstrom, g., nettles, m. & abers, g.a. 2003: glacial earthquakes. science 302, 622–624. ekstrom, g., nettles, m. & tsai, v.c. 2006: seasonality and increasing frequency of greenland glacial earthquakes. science 311, 1756–1758. gregersen, s. 1989: the seismicity of greenland. in: gregersen, s. & basham, p.w. (eds): earthquakes at north-atlantic passive margins: neotectonics and postglacial rebound, 345–353. berlin: springer. harrison, w., echelmeyer, k. a. & engelhardt, h. 1993: short-period observations of speed, strain and seismicity on ice stream b, antarctica. journal of glaciology 39, 463–470. johnston, a.c. 1987: suppression of earthquakes by large continental ice sheets. nature 330, 467–469. kumar, p. et al. 2005: the lithosphere–asthenosphere boundary in the north-west atlantic region. earth and planetary science letters 236, 249–257. kumar, p., kind, r., priestley, k. & dahl-jensen, t. 2007: crustal structure of iceland and greenland from receiver function studies. journal of geophysical research 112, b03301.1–b03301.19. larsen, t.b., dahl-jensen, t., voss, p., jørgensen, t.m., gregersen, s. & rasmussen, h.p. 2006: earthquake seismology in greenland – improved data with multiple applications. geological survey of denmark and greenland bulletin 10, 57–60. lund, b. & näslund, j.-o. 2009: glacial isostatic adjustment: implications for glacially induced faulting and nuclear waste repositories. in: connor, c.b., chapman, n.a. & connor, l.j. (eds): volcanic and tectonic hazard assessment for nuclear facilities, 142–155. cambridge: cambridge university press. nettles, m. et al. 2008: step-wise changes in glacier flow speed coincide with calving and glacial earthquakes at helheim glacier, greenland. geophysical research letters 35, l24503, doi:10.1029/2008gl036127. o’neel, s., mcnamara, d., marshall, h. & pfeffer, t. 2006: seismic evidence for time variation of mechanical failure associated with iceberg calving at columbia glacier, ak. eos transactions, american geophysical union, 87(52), 1719 only. rial, j.a., tang, c. & steffen, k. 2009: glacial rumblings from jakobshavn ice stream, greenland. journal of glaciology 55, 389–399. stewart, i.s., sauber, j. & rose, j. 2000: glacio-seismotectonics: ice sheets, crustal deformation and seismicity. quaternary science reviews 19, 1367–1389. ucisik, n., gudmundsson, o., hanka, w., dahl-jensen, t., mosegaard, k. & priestley, k. 2008: variations of shear-wave splitting in greenland: mantle anisotropy and possible impact of the iceland plume. tectonophysics 462, 137–148. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: tdj@geus.dk geological survey of denmark and greenland bulletin 25, 2011 + 1 plate 11 geological survey of denmark and greenland bulletin 25 • 2011 upper cretaceous chalk facies and depositional history recorded in the mona-1 core, mona ridge, danish north sea kresten anderskouv and finn surlyk geological survey of denmark and greenland danish ministry of climate, energy and building 22 geological survey of denmark and greenland bulletin 25 keywords upper cretaceous chalk, danish north sea, mona-1, facies analysis cover detail of a shear-deformed, matrix-supported chalk conglomerate representing a deformed debris-flow deposit (mona-1 polished core, 10890 ft). such chalk deposits are typical of the upper cretaceous succession in the mona-1 well and exemplify two important processes of redeposition: fluid-like deformation during debris flow and plastic deformation during slumping or the final stages of debris-flow deposition. photograph courtesy of wintershall noordzee b.v. chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, department of geoscience, aarhus university; minik rosing, geological museum, university of copenhagen; finn surlyk, department of geography and geology, university of copenhagen scientific editor of this volume: jon r. ineson editorial secretaries: jane holst and esben w. glendal referees: frans van buchem (dk) and maurice e. tucker (uk) illustrations: jette halskov, britta munch and christian hagen digital photographic work: benny m. schark graphic production: kristian a. rasmussen printers: rosendahls . schultz grafisk a/s, albertslund, denmark manuscript received: 6 december 2010 final version approved: 19 october 2011 printed: 29 december 2011 issn 1604-8156 isbn 978-87-7871-327-8 citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 25, 60 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk or at www.geus.dk/publications/bull © de nationale geologiske undersøgelser for danmark og grønland (geus), 2011 for the full text of the geus copyright clause, please refer to www.geus.dk/publications/bull 33 contents abstract . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . setting and stratigraphy . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . materials and methods . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . facies analysis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . definitions and terminology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . facies 1: bioturbated chalk . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . facies 2: bioturbated marly chalk . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . facies 3: laminated chalk . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . facies 4: laminated marl . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . facies 5: graded silty chalk . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . facies 6: graded silty chalk with packstone lamination . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . facies 7: graded packstone and wackestone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . facies 8: non-graded packstone and wackestone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . facies 9: structureless chalk . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . facies 10: structureless matrix-supported chalk conglomerate . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . facies 11: clast-supported chalk conglomerate . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . facies 12: shear-deformed bioturbated chalk . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . facies 13: shear-deformed matrix-supported chalk conglomerate . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . facies 14: shear-banded chalk . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . slumps . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . genetically related facies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . facies associations and stratigraphic development . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . facies associations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . facies association 1 (pelagic chalk) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . facies association 2 (turbiditic chalk) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . facies association 3 (mass-transport chalk) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . stratigraphic development . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . basin development and depositional history . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . stacking pattern and predictability . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . references . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5 6 8 10 10 10 12 14 16 17 18 20 22 23 23 24 29 32 34 39 42 44 46 46 46 46 46 48 53 53 56 57 57 58 44 55 anderskouv, k. & surlyk, f. 2011: upper cretaceous chalk facies and depositional history recorded in the mona-1 core, mona ridge, danish north sea. geological survey of denmark and greenland bulletin 25, 60 pp. the 331 m long core from the mona-1 well in the danish north sea spans almost the entire upper cretaceous chalk group but only about 10% of late cretaceous time is represented. the succession comprises 14 facies representing pelagic deposition, turbidity flow, and mass-transport processes, including mudflow, debris flow, and slumping. pelagic deposits vary mainly in terms of the concentration of siliciclastic material, the trace-fossil assemblage, and the presence or absence of primary sedimentary structures. pelagic sedimentation was probably punctuated by the deposition of thin turbidites, and the resultant deposits were thoroughly bioturbated if deposited during normal oxygenation at the sea floor. periodic benthic dysoxia resulted in the preservation of primary structures, as represented by laminated chalk which consists of thin pelagic laminae alternating with thin turbidites. in addition to the thin turbidites in the laminated chalk, four different turbidite facies are interpreted as representing highto low-energy flows. clast-supported chalk conglomerates have previously not been differentiated from other turbidites, but are here interpreted to be directly related to the down-slope evolution of debris flows. debris flows are represented by matrix-supported conglomerates, which form one of the most common facies in the succession. high-concentration, gravity-driven suspension flows passed into dilute visco-plastic flows during the final stages of deposition and resulted in the deposition of structureless chalks. limited shear deformation produced distinct quasi-facies from which the precursor facies can be deduced, whereas intense or continued shear deformation produced a shear-banded quasi-facies from which the precursor facies cannot be deduced in all cases. a series of major slump packages (14–18 in total) are interpreted, forming over 40% of the succession; debrites appear to be the most common precursor facies involved in slumping. the vertical succession of facies records an earliest cenomanian facies shift from dominantly siliciclastic to chalk deposition. the cenomanian – late campanian period was dominated by erosion or sediment by-pass with minor associated mass-transport deposits preserved. basin filling by pelagites and turbidites prevailed in the late campanian, whereas maastrichtian pelagic deposition was interrupted by increasingly frequent and voluminous mass-transport events. authors’ address department of geography and geology, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ka@geo.ku.dk, finns@geo.ku.dk abstract 66 introduction the upper cretaceous chalk group has long been the subject of research concentrating on stratigraphy, palaeontology and geochemistry. few studies have addressed the physical processes of deposition in detail, with some notable exceptions (e.g. voigt 1962, 1977; voigt & häntzschel 1964; steinich 1967, 1972; kennedy & juignet 1974; kennedy & garrison 1975; gale 1980; hardman & kennedy 1980; noe-nygaard & surlyk 1985; quine & bosence 1991; scholle et al. 1998; damholt & surlyk 2004; anderskouv et al. 2007), and almost no empirical data on the depositional processes of chalk ooze exist. as a result, the literature contains few examples of depositional models for chalk and none of those published have been demonstrated to be successfully predictive, with respect to identifying the occurrence of a particular facies based on the known occurrence of other facies. vertical successions of chalk facies in the chalk group of the north sea have been considered non-ordered, unlike the commonly systematic and predictable stacking patterns of deep-water sili-ciclastic successions (e.g. surlyk et al. 2003). although this is mainly due to the line-sourced nature of most carbonate redeposition systems and the fact that most redeposition processes may be initiated at any depth on the slope (e.g. mullins & cook 1986), it probably also reflects a relatively limited understanding of the lateral associations and gradations between processes and facies in the chalk sea. descriptive classification schemes for chalk facies or lithologies have been introduced (hancock 1975; siemers et al. 1994; crabtree et al. 1996). the understanding of depositional systems can be used to predict the distribution of facies or lithologies, a task that is interpretive by nature. objective description is necessary to document and communicate an interpretation, but it has no predictive power in itself. a variety of interpretive facies and conceptual depositional models for the north sea chalk group have also been proposed (nygaard et al. 1983; brewster et al. 1986; kennedy 1987a, b; nielsen et al. 1990; sikora et al. 1998; bramwell et al. 1999). such models may prove predictive when based on a solid foundation of empirical data and/or nume rous well-documented cases. the purpose of the present paper is to contribute to such a foundation by presenting a detailed study of a long and varied succession. the analysis builds upon all the papers cited facing page: fig. 2. a: main structural elements in the north sea central graben. modified from surlyk et al. (2003). dk: denmark. ger.: germany. neth.: netherlands. uk: united kingdom. b: structural elements and chalk group isochore contours in metres in the mona ridge area. the shaded area, where the chalk group is less than 600 m thick, broadly defines the position of the mona ridge. the dashed line indicates the location of the seismic line in fig. 3. modified from britze et al. (1995). fig. 1. palaeogeography of north-west europe during the cenomanian–turonian (a), and the coniacian–maastrichtian (b). modified from ziegler (1990). north sea marine depositslate cretaceous landmasses north sea cenomanian–turonian coniacian–maastrichtian b a 200 km n 200 km n 77 jeppe basin ringkøbing– fyn high mandal high lindesnes ridge norwegian– danish basin tail end graben inge high 2°e 4°e 57°n 56°n søgne basin coffee soil fault auk ridge josephine high forties– montrose high west central graben east central graben arne ridge a uk dk ger.neth. norway b mid north sea high mona ridge mandal high karl basin mona-1 karl-1 jeppe basin norway denmark 700 800 600 500 900 600 1000 1300 600 500 700 60 0 1 km 0 20 40 km fault mona ridge 88 above although the facies descriptions do not rely on a classification scheme and interpretations do not follow a specific depositional model. nor is the applied facies subdivision presumed to constitute a universally employable facies scheme. the mona-1 core is ideal for the purpose of the study as it covers a 303 m long chalk succession ranging from the base upper cretaceous almost to the cretaceous–tertiary boundary. a wide range of chalk facies are represented in the succession, which thus provides an excellent opportunity for analysing the vertical development of chalk facies in the upper cretaceous part of the chalk group in the north sea. the aim of this study is to present detailed descriptions and interpretations of the chalk facies present, and to discuss facies associations and the general late cretaceous depositional history in the mona ridge area. setting and stratigraphy during the late cretaceous, much of north-west europe was covered by an extensive epeiric sea that was characterised by low terrigenous input and deposition of chalk (håkansson et al. 1974; hancock 1976; surlyk 1997; surlyk et al. 2003). the central north sea area constituted a relatively deep part of the chalk sea, flanked to the east by the wide pelagic carbonate ramp and basin of the danish basin, the siliciclastic-dominated seas towards the north, the shallower seas around the present british isles to the west, and the paris basin to the south (fig. 1). the central graben is a generally nnw-trending intracratonic basin in the north sea, delimited by the coffee soil fault to the east and the mid north sea high to the west. it was formed by permian and triassic rifting, middle jurassic thermal uplift and late jurassic rifting along nto ne-trending normal faults (e.g. zanella & coward 2003). thermal subsidence characterised the late cretaceous, punctuated by inversion pulses (cartwright 1989; vejbæk & andersen 2002). the mona ridge is an elongate nw– se-trending structural high in the northernmost part of the danish north sea near the eastern limit of block fig. 3. seismic section from the eastern part of the jeppe basin (left) to the western margin of the mandal high (right); for location of seismic line, see fig. 2. interpretation courtesy of f.c. jakobsen (geus). mona-1 karl-1 3000 3500 4000 tw ow ay tr av el ti m e (m se c) w e top chalk top tor intra-tor unconformity top hod base chalk base cretaceous 1 km 99 fig. 4. lithostratigraphic subdivision of the chalk group in the north sea. modified from surlyk et al. (2003). l: lower. m: middle. u: upper. fig. 5. biostratigraphy of the mona-1 core based on calcareous nannofossils. modified from bailey et al. (1999). grey parts denote lack of nannofossil data due to gaps between samples or poor preservation. alb.: albian. cenoman.: cenomanian. l.: lower. cr.: cretaceous.central north sea stratigraphy system series stage pa lae og en e pa leo ce ne danian c ro m er kn ol l g ro up ro ga lan d g ro up ekofisk formation tor formation hod formation blodøks formation hidra formation ch alk g ro up cr et ac eo us up pe r lo we r al bia n up pe r up pe r up pe r lo we r lo we r l m m idd le m m m l l l u u u u ce no m an ian tu ro nia n ca m pa nia n m aa str ich tia n co nia cia n sa nto nia n 11300 11250 11200 11150 11050 11100 11000 10950 10900 10850 10800 10750 10700 10650 10600 10500 10550 10450 10400 10350 3150 3160 3170 3180 3190 3200 3210 3220 3230 3240 3250 3260 3270 3280 3290 3300 3310 3320 3330 3340 3350 3360 3370 3380 3390 3400 3410 3420 3430 3440 3450 d ep th (f t) d ep th (m ) se rie s st ag e n an no pl an kt on z on e up pe r m aa st ric ht ian lo w er m aa st ric ht ian up pe r c am pa ni an c en om an . uc20 uc19 uc18 uc16 bc27 u pp er c re ta ce ou s uc10 uc3* uc11 ** *** al b. l. c r. * uc3 or older ** middle–upper coniacian *** upper coniacian – lower santonian 1010 21 (fig. 2). it is flanked to the sw by the wide, nw– se-trending jeppe basin and to the ne by the narrow and deep karl basin (fig. 3). the karl basin is delimited by the mona ridge and the mandal high further to the ne. the mona-1 well is situated just east of the crest of the mona ridge adjacent to the karl basin. the most logical and coherent lithostratigraphic scheme for the upper cretaceous in the north sea appears to be that followed by surlyk et al. (2003). this is essentially the scheme of deegan & scull (1977) with the modifications of isakson & tonstad (1989), but retaining the original distinction between the dominantly siliciclastic shetland group in the northern north sea and the carbonate-dominated chalk group in the southern and central north sea (fig. 4). few biostratigraphical studies have been undertaken on the upper cretaceous succession in mona-1; the nannofossil zonation established for the well by bailey et al. (1999) is used here (fig. 5). materials and methods the mona-1 core is 331 m long and includes the uppermost part of the lower cretaceous cromer knoll group and 303 m of the upper cretaceous chalk group. core recovery was generally very high, except for a c. 12 m section in the upper part of the core. the core condition is excellent in the lower half of the upper cretaceous, whereas the porous upper part is in moderately good condition but very fragile and friable. the studied succession extends from the uppermost albian to less than 10 m below the cretaceous–tertiary boundary at the top of the core. a detailed core description was made at a scale of 1:5 or 1:15 in order to record the subtle sedimentary variations. a number of polished slabs were made, treated with light oil, and scanned at a resolution of 1200 dpi. this provides the equivalent of a 10-fold magnification at optimal resolution and allows for subsequent digital image processing such as contrast enhancement. all presented sedimentological logs are drawn in a naturalistic style, as close to the actual appearance of the core as possible. trace fossils are identified to generic level. twenty-five thin sections were produced for petrographic analyses. porosity and permeability data were obtained from the conventional core analyses performed in 1983. the original plug and drilling data are reported in imperial units, which were also used during core description to minimise the risk of error. depths were subsequently converted to metric units and both types of units are reported throughout the manuscript. facies analysis definitions and terminology descriptions and interpretations of facies are summarised in table 1. some of the facies described below would probably be recognised as representing a continuum if studied in outcrop; such gradual change is very difficult to determine in core. a large number of facies was deliberately distinguished to prevent loss of information, as subtle differences have a potentially significant impact on the interpretation of the succes1111 sion (table 1). texture is classified according to the modification of dunham (1962) by embry & klovan (1971), with the further modification that mud is defined according to the wentworth scale so that grains are particles larger than 63 μm. the terms ‘silty’ and ‘sandy’ are used solely to describe grain sizes and do not imply siliciclastic content. the terms ‘plastic’ and ‘visco-plastic’ are used according to mechanical definitions (middleton & wilcock 1994), i.e. ‘plastic’ is not synonymous with ‘ductile’ or ‘soft’. the term ‘bioturbated’ is distinguished from the term ‘burrowed’ in this study. ‘bioturbated’ is used here for extensively burrowed and biomottled deposits with no preserved primary sedimentary structures, such as lamination. facies description interpretation 1 bioturbated chalk completely bioturbated chalk with no preserved pelagite primary sedimentary structures. 2 bioturbated marly chalk completely bioturbated marly chalk with pelagite no preserved primary sedimentary structures. 3 laminated chalk chalk with vague and indistinct lamination seen pelagites alternating with as subtle colour differences. low-density turbidites and fall-out deposits from clouds of resuspended material 4 laminated marl carbonate-rich siliciclastic mudstone with irregular, pelagite/hemipelagite wispy and laterally discontinuous lamination. 5 graded silty chalk sharply based chalk with a relatively high turbidite concentration of silt-grade particles. burrowed and rarely laminated. 6 graded silty chalk with sharply based graded chalk with a relatively high turbidite packstone lamination concentration of silt-grade particles and fine sand to silt laminae. 7 graded packstone and sharply based fine-sand packstone, which grades turbidite wackestone upwards into a fine sandy chalk wackestone. 8 non-graded packstone fine-sand packstone or fine sandy wackestone with turbidite and wackestone a high density of burrows. 9 structureless chalk structureless chalk, which may or may not be mudflow deposit burrowed to a limited extent. 10 structureless matrixstructureless chalk or wackestone with randomly debrite supported chalk distributed and oriented matrix-supported fine conglomerate sand to boulder-sized clasts, which may show coarse-tail grading. 11 clast-supported chalk pebbly grainstone to pebble conglomerates high-density turbidite, conglomerate composed of chalk clasts and subordinate shell debris. derived from debris flow 12 shear-deformed deformed chalk in which the bioturbated fabric limited plastic deformation of bioturbated chalk is still recognisable. bioturbated chalk 13 shear-deformed matrixmatrix-supported chalk conglomerate with deformed debrite supported chalk deformation structures. conglomerate 14 shear-banded chalk matrix-supported chalk conglomerate with irregular intensely shear-deformed colour banding. chalk table 1. facies summary 1212 ‘burrowed’ is used for deposits that have been burrowed, but not to the extent that primary sedimentary structures have been totally obliterated. fourteen facies are recognised, but facies 12–14 are not facies in a classical sense, as they are defined by structures caused by penecontemporaneous deformation. however, they are defined as facies here because the occurrence of soft sediment deformation is critical for the interpretation of the succession. also, it is in some cases impossible to determine whether the penecontemporaneous deformation was purely post-depositional or part of the depositional process. the primary lithological discriminator is the distinction between chalk, marly chalk, and marl (plate 1). chalk is defined as a rock or sediment that is compositionally dominated by calcareous nannofossils (scholle 1977). in the studied succession, the chalk is a carbonate mudstone, and rarely a wackestone, and the main constituents of the carbonate mud are calcareous nannofossils with subordinate calcispheres and small foraminifers. sand-sized bioclasts are mainly foraminifers, inoceramid shell fragments, and small fragments of bivalves, echinoderms, and other invertebrate fossils. the term marly chalk is used for chalk with a sufficiently high clay content to be recognised by its colour with the naked eye; the term marl is used for carbonate-bearing siliciclastic mudstones. facies 1: bioturbated chalk description. this facies is characterised by its lithological composition and a completely bioturbated fabric with no preserved primary sedimentary structures (figs 6, 7). the most common recognisable burrows are large and small chondrites and planolites and, less common, teichichnus and taenidium. zoophycos is largely restricted to the 3288–3284 m (10 788 ft – 10 774 ft) interval, and possible thalassinoides burrows mainly occur in the upper half of the core. the facies most commonly has a mudstone texture and may contain few scattered sand-sized bioclasts. in rare cases, local wackestone texture occurs due to the concentration of bioclasts. lower boundaries are commonly gradational to the underlying facies, whereas upper boundaries are most commonly sharp. the facies constitutes a little more than one third of the entire succession and individual facies units are up to 6.3 m thick. it occurs throughout the succession, but it is most important in the maastrichtian part. it is commonly interbedded with facies 5–8 and less commonly with facies 3 and 4. intrapelagic erosion surfaces are important features intimately associated with facies 1. they are sharp surfaces occurring within bioturbated chalk intervals, typically expressed by a subtle but abrupt change in colour. the surfaces appear to be smooth but may show a slightly wavy topography. they truncate burrows of the underlying unit, but may be penetrated by burrows descending from the overlying chalk. there are no significant changes in composition or texture associated with the surfaces. a total of 27 intrapelagic erosion surfaces are recognised, restricted to the lower half of the studied succession and strongly concentrated in the 3360–3325 m (11 030 ft – 10 910 ft) interval. they are most common within pelagic units that include relatively few, thin and fine-grained interbeds belonging to facies 5–8. interpretation. the complete lack of primary structures and the thoroughly bioturbated fabric indicate fully oxygenated benthic conditions and relatively slow accumulation. these conditions allowed the infauna to continuously burrow the sea-floor sediments and destroy all primary structures. the range of trace fossil tiers and the rarity of benthic fossils indicate a relatively rich infauna and a restricted shelly epifauna, which suggests a high nutrient flux to the sea floor (dayton & oliver 1977; bambach 1993; mckinney & hageman 2006). the original depositional process cannot be determined due to the destruction of primary structures, but the slow accumulation indicates deposition mainly by pelagic fall-out although a contribution from thin event beds is likely (see facies 3). the lack of mineralisation or lags in association with intrapelagic erosion surfaces indicates that they do not represent significant time intervals, but rather appear to represent short perturbations of the pelagic deposition by sediment gravity flows or bottom currents; the latter are increasingly being recognised as an important influence on relatively deep-water chalk deposition (lykke-andersen & surlyk 2004; esmerode et al. 2007, 2008; surlyk & lykke-andersen 2007; surlyk et al. 2008; esmerode & surlyk 2009). the lack of lags and characteristic event deposits suggests that the eroded material was transported away from the site, rather than being resuspended locally and producing a postevent, fall-out deposit. the amount of erosion is difficult to assess in a core, but the fact that the surfaces are preserved implies that the soft uppermost part of the 1313 sediment column, where the most efficient burrowing organisms would have prevailed, was removed by erosion. deep burrowers adapted to firm, somewhat compacted sediment (e.g. chondrites) were not excluded, however, indicating that erosion probably proceeded to a depth of a few tens of centimetres. 1 cm fig. 6. typical appearance of bioturbated chalk (facies 1; 3360.02–3359.87 m, see plate 1). note the mottled appearance due to the numerous superimposed vaguely defined colour variations. 1414 facies 2: bioturbated marly chalk description. this facies comprises completely bioturbated chalk with a siliciclastic mud content large enough to be detected by its colour with the naked eye (figs 8, 9). the ichnofabric is dominated by small chondrites, zoophycos, planolites, with less common large chondrites and taenidium. the facies only occurs in a 3.15 m thick interval in the cenomanian part of the core, where it transitionally overlies laminated marl (facies 4). interpretation. the thorough bioturbation indicates relatively slow pelagic carbonate and hemipelagic clay accumulation under well-oxygenated benthic conditions. the prevalence of zoophycos may indicate relatively stable conditions and slow continuous deposition in deep water (ekdale & bromley 1984). mud sand grf m c pebbles 10705 10708 10711 10714 10717 10702 10704 10707 10710 10713 10716 10703 ftm 10706 10709 10712 10715 3262 3263 3264 3265 3266 fig. 7. typical bioturbated succession showing random distribution of bioclasts and recognisable trace fossils, mostly planolites, chondrites, some teichichnus, and one possible thalassinoides. burrows shell fragments intrapelagic erosion surfaces biomottling 1515 fig. 9. log example of bioturbated marly chalk (facies 2). note the intensely bioturbated fabric dominated by small chondrites, zoophycos and planolites. 1 cm fig. 8. bioturbated marly chalk (facies 2; 3441.50–3441.24 m, see plate 1) with a mottled appearance and prominent zoophycos (arrows). 11295 ft 11296 11297 3443 3442.6 m zoophycos on biomottled background marly interval with burrows 1616 facies 3: laminated chalk description. this facies is characterised by vague and indistinct lamination expressed as subtle colour differences (figs 10, 11). the laminae are 1–10 mm thick, planar, commonly laterally discontinuous, and lowangle cross-cutting relations between laminae have been observed. the laminae are not visible under the microscope, as there are no obvious differences in grain size, composition or texture. the cause for lamination must thus be variations on the nannofossil scale. in rare cases, the laminae appear to be inversely graded, and small-scale dish structures also occur. it is common to find diagenetic lamination induced by slight pressure dissolution in association with true lamination. true lamination is distinguished from diagenetic lamination by being cut by burrows and lacking signs of pressure solution and concentrations of dissolution residues. some units are devoid of burrows, but burrows commonly occur in small numbers. chondrites and small planolites are by far the most common, but teichichnus has also been observed. the facies makes up a total of 5.5 m distributed over 37 occurrences throughout the core and individual units are 2–150 cm thick. eight units are thicker than 30 cm, 12 are 5–30 cm thick, and 17 are thinner than 5 cm. most of the units thicker than 30 cm are overlain by redeposited facies, and a little more than half of the 5–30 cm thick units are overlain by redeposited facies, whereas most of the thin units are overlain by bioturbated facies. interpretation. the lamination was probably formed by deposition from low-density turbidity currents alternating with pelagic fall-out and settling from clouds of suspended material (damholt & surlyk 2004). the suspended material may have been resuspended by currents or waves or lofted from sediment gravity flows. all these processes produced thin laminae or surfaces which would rapidly have been destroyed by bioturbation under oxygenated benthic conditions. the preservation of lamination and the lack of bioturbation were thus interpreted by damholt & surlyk (2004) to have been the result of dysoxia at the sea floor during deposition. other authors have interpreted such laminated chalk beds to have formed by rapid deposition, the high sedimentation rate precluding thorough bioturbation (scholle et al. 1998). the depositional process and the cause for the lamination according to this latter model are unclear, however, and the interpretation of damholt & surlyk (2004) is thus preferred here. the fall velocity of coccoliths and microfossils, such as calcispheres and foraminifers, differ by two orders of magnitude (mccave 2008), and deposition from dilute suspension clouds or turbidity flows may therefore be expected to have produced clear distribution grading, a feature not observed in this laminated chalk facies of the mona-1 core (see also damholt & surlyk 2004). there may be several explanations for the lack of well-defined distribution grading. lofted flows and nepheloid layers would probably consist of only very fine grains, thus having a too limited grain-size range to produce visible 1 cm fig. 10. laminated chalk (facies 3; 3386.01–3385.93 m, see plate 1). the image shows both primary (large arrows) and diagenetically enhanced lamination (small arrows); note that primary lamination passes laterally into diagenetic lamination in some cases. burrows occur throughout. 1717 distribution grading. indeed, the finest parts of other chalk turbidites studied here, consisting largely of coccoliths and calcispheres, typically do not show marked distribution grading. it must be borne in mind that although chalk ooze is commonly referred to as ‘completely non-cohesive’ due to the lack of clay minerals, almost nothing is known about the physical properties of the ooze in its original form. van der waals forces, organic material such as extracellular polymeric substances, grain morphology and non-preserved spiculate material are all potentially highly influential factors in this respect. thus, particle aggregation may have been a significant factor in such fine-grained suspensions, and this in turn may have significantly influenced the resultant grain-size distribution after the removal of labile material. the close association between the thickness of laminated units and the type of overlying deposits suggests a causal link. if deposition took place during dysoxia and was followed by similar deposition under oxygenated conditions, burrowing organisms would penetrate downwards into the laminated sediment, thereby destroying the lamination and reducing the thickness of the laminated unit. however, the full thickness of the laminated unit would be preserved if the unit was capped by redeposited chalk. this may explain why thick laminated units are commonly overlain by redeposited facies, whereas thin laminated units are typically overlain by bioturbated facies. facies 4: laminated marl description. this facies comprises carbonate-rich siliciclastic mudstone with irregular, wispy, and laterally discontinuous lamination (fig. 11). it makes up a total of 3.60 m of the studied succession and occurs at four different levels and in three somewhat different fashions. the lowermost part of the studied succession is a laminated marlstone with abundant chondrites and common planolites. it passes upward into bioturbated marly chalk at the albian–cenomanian transition. the facies also occurs at 3423.7–3422.7 m (11 232 ft 8 in. – 11 229 ft 4 in.) where it transitionally overlies a 46 cm thick turbidite (facies 6). in this unit, the lamination is less well preserved due to the abundance of shallow tier planolites and taenidium. in two other cases in the 3411.6–3408.0 m (11 193 ft – 11 181 ft) interval, 11181 11182 11183 11184 11185 3408.5 3408 m ft 3409 lamination marl burrows fig. 11. log example of laminated chalk (facies 3, see plate 1) alternating with thin beds of laminated marl (facies 4). note the irregularity of the lamination. the marl beds are invariably associated with pressure dissolution. 1818 laminated marl occurs as numerous diagenetically enhanced layers interbedded with laminated chalk. the marl layers are invariably associated with dissolution features, and they are generally c. 1 cm thick. interpretation. the units belonging to this facies may have polygenetic origins. the lowermost example is interpreted to represent hemipelagic and some pelagic deposition during oxygen-deficient benthic conditions. the occurrence at 3423.7–3422.7 m (11 232 ft 8 in. – 11 229 ft 4 in.) appears to represent deposition from dilute mud suspension following deposition of the underlying turbidite. the thin dissolution marl laminae associated with laminated chalk may represent primary marl layers, or they may simply be pressure-solution products of originally slightly argillaceous chalk. facies 5: graded silty chalk description. this facies is conspicuous due to its pale colour, in contrast to the dominantly light grey colour of the surrounding chalk (fig. 12). beds are 1–22 cm thick and invariably occur within bioturbated intervals; they are typically heavily burrowed and their primary structures are only partially preserved. the facies has a relatively high concentration of silt-grade and more rarely fine sand-grade particles, mainly calcispheres (fig. 13a). where preserved, beds show a subtle overall grading with 0.5 mm thick silt laminae. the laminae have high concentrations of calcispheres and rarer small foraminifers, both of which are typically filled with pyrite or calcite, giving the laminae a black coloration (figs 12, 13a). lower boundaries are sharp or burrowed, and upper boundaries are invariably burrowed. there are 35 beds of facies 5 in the core, principally in the 3378.6–3324.1 m (11 084 ft 9 in. –10 905 ft 10 in.) interval. 1 cm fig. 12. graded silty chalk (facies 5; 3369.44–3369.36 m, see plate 1), which has a conspicuous pale colour and is burrowed. the black parts represent remnants of silty lamination heavily cemented by pyrite (arrow). 1919 a b 250 µm 100 µm interpretation. the sharp base, overall grading and graded silty laminae all indicate that the facies represents silty mud-grade carbonate turbidites, which due to their limited thickness were heavily burrowed subsequent to deposition. in terms of standard turbidite models, the facies is analogous to the lower part of the e division of bouma (1962), the upper part of e1 of piper (1978), or the t2–t4 divisions of stow & shanmugam (1980). fig. 13. photomicrographs (plane polarised transmitted light) of thin sections. a: silty laminae in graded silty chalk (facies 5; 3369.40 m, see plate 1). note absence of intraparticle porosity due to pyrite and calcite cement. b: silt to fine sand lamina in graded packstone and wackestone (facies 7; 3284.36 m, see plate 1) showing welldeveloped intraparticle cement. 2020 facies 6: graded silty chalk with packstone lamination description. facies 6 is composed of chalk, and silt to fine sand-grade pelagic bioclasts and fragments of benthic invertebrate fossils. the 10–45 cm thick beds all have sharp and commonly loaded bases (figs 14, 15), which truncate burrows of underlying facies. the base is typically overlain by a graded and in some cases laminated sandy interval a few centimetres thick (fig. 16). the basal interval grades upwards into a graded silty chalk interval with discrete fine sand to silt laminae that show upwards decreasing, lateral continuity, thickness and regularity. the laminae are reduced to thin, indistinct silty lenses towards the top of the beds. the graded silty chalk interval makes up the greater part of the beds. one ripple form set and several examples of small-scale dish structures have been recorded. the top of the beds may comprise a relatively thin structureless chalk interval grading into bioturbated chalk, or the graded chalk may be overlain by a thick laminated chalk or marl that is burrowed at the top (facies 3 and 4). the beds are not burrowed, except at the top when overlain by bioturbated chalk. there are 13 beds referred to facies 6, which make up a total of 3.18 m. all beds occur in the 3411.6– 3398.1 m (11 192 ft 9 in. – 11 148 ft 9 in.) interval, except for one at 3424.2 m (11  234 ft 2 in.). thus, the facies occurs in association with facies 7 and facies 1, and three of the beds are overlain by the laminated facies 3 and facies 4. interpretation. the sharp erosional base, overall normal grading and succession of structures indicate that the facies represents fine sand to mud turbidites. they essentially correspond to the c–e divisions of bouma (1962), e1–e3 of piper (1978), and t0–t8 of stow & shanmugam (1980) although facies 6 is starved of sand-sized particles compared to siliciclastic turbidites, which formed the basis for the standard facies models mentioned above. the siliciclastic examples are distal fig. 14. graded silty chalk with packstone lamination (facies 6; 3424.21–3423.84 m, see plate 1). the lower part (bed base is indicated) contains several barely visible contorted sand-grade laminae, whereas lamination is more clearly visible in the upper part of the sample due to colour variations caused by porosity differences. 2121 shelfor delta-derived, outer-fan and channel-levee deposits, whereas facies 6 was derived from pelagic chalk with a low content of sand-sized carbonate particles. fig. 15. log example of a graded silty chalk bed with packstone lamination (facies 6, see plate 1). bed base at 3424.13 m. fig. 16. detail (a) and photomicrograph (b) of thin section of a lamina in facies 6 (3407.13 m, see plate 1). note the normal grading in b. the dashed outline in a indicates the location of the thin section in b. mud sand gr f m c 11232 11233 11234 3423.5 m ft 3424 burrows lamination bioclastic sand lamination 2 mm 500 µm a b 2222 facies 7: graded packstone and wackestone description. this facies may be regarded as equivalent to the lower, graded sandy interval of facies 6 being characterised by sharply based packstone which grades upwards into chalk wackestone (fig. 17). it consists of silt to fine-grained pelagic carbonate sand and pelagic mud, together with benthic bioclastic debris which is mostly up to fine sand grade although slightly coarser fig. 17. graded packstone and wackestone (facies 7). a: two thin sandy beds of facies 7 units (3376.04–3375.86 m, see plate 1) characterised by sharp bases, normal grading and top–down burrowing. b: a facies 7 sample (3401.97–3401.87 m, see plate 1) with a relatively thick, laminated sand-grade base. c: facies 7 example (3284.40–3284.32 m, see plate 1) with heavy pyrite staining. sand-sized bioclasts are commonly filled by pyrite cement in all three examples, but the cementation is especially conspicuous in c. 2323 bioclasts do occur. faint parallel lamination is relatively common and ripple cross-lamination occurs in one unit. burrows are common in the top of the beds but rare in the basal part. the lower boundaries are invariably sharp and truncate underlying structures, whereas upper boundaries are gradational unless truncated by another bed belonging to facies 5–8. facies 7 is commonly partially stained black by pyrite and thin sections invariably show well-developed intraclastic cementation (fig. 13b). facies 7 almost always overlies, and is overlain by, bioturbated chalk of facies 1. there are 40 facies 7 beds making up a total of 2.62 m in the core, with individual beds being 1–18 cm thick. interpretation. facies 7 beds are interpreted as turbidites on the basis of the sharp, erosional lower boundary, normal grading, horizontal lamination, rare ripple cross-lamination, and top-down burrowing. the facies belongs to the c division of bouma (1962), the basal part of e1 of piper (1978), or the t0 division of stow & shanmugam (1980). in many cases, the facies resembles the basal part of facies 6, the main difference being that facies 7 is slightly thicker and sandier than the basal part of facies 6, and that it lacks the relatively thick muddy interval with sandy laminae that dominates facies 6. facies 8: non-graded packstone and wackestone description. this facies is represented by thin beds of fine sand packstone and more rarely wackestone. the composition is identical to that of facies 7. the facies is normally a few centimetres thick, shows a high density of burrows, and lower and upper boundaries are transitional to fully bioturbated chalk (facies 1), except in two cases where it is associated with redeposited facies and where facies 8 shows no burrows and is sharply bounded by thin zones of intense shear deformation. there are 26 occurrences of the facies making a total of 80 cm, almost all of which are found in the pelagite/turbidite dominated interval in the lower half of the core (3411.6–3324.1 m, 11 192 ft 9 in. – 10 905 ft 10 in.). interpretation. facies 8 is identical to facies 7 in terms of composition and grain size, but lacks grading, sharp boundaries and sedimentary structures. it is interpreted to represent thin, fine sand-grade turbidites, equivalent to facies 7, which were homogenised by bioturbation, or in rare cases by shear deformation. burrowing generally affected only the top part of facies 7, whereas the generally thin facies 8 units suffered complete biogenic reworking. the bioturbation destroyed grading and other structures, blurring the sharp boundaries and, in some cases, resulting in mixing of the turbidite sand and chalk ooze and the creation of a wackestone texture. more rarely, the homogenisation was caused by shear deformation rather than bioturbation. the facies belongs to the c division of bouma (1962), the basal part of e1 of piper (1978), or the t0 division of stow & shanmugam (1980). facies 9: structureless chalk description. this facies is a structureless chalk, which may be burrowed to a limited extent (fig. 18). it may contain fine sand-grade bioclasts and rare randomly distributed and oriented benthic fossil fragments. in some cases, the facies shows subtle signs of deformation represented by vague, folded colour-bands. it does not show the mottled appearance that characterises the bioturbated chalk facies. burrows are typically few and concentrated at the top of individual beds. the most common burrows are small chondrites, whereas planolites, taenidium, and teichichnus occur in subordinate numbers. trichichnus, which is very rare in the other facies, is relatively common. some facies 9 intervals show internal burrowed surfaces from which burrows penetrate down into the underlying bed, whereas the overlying beds are devoid of burrows, indicating that the interval is amalgamated and composed of several individual units. lower boundaries are typically transitional over a short vertical distance and therefore appear abrupt. there are 28 units of the facies, which mainly occur in intervals characterised by extensive redeposition. individual units are 2–815 cm thick, and make up a total of 23 m. most units are less than 30 cm thick and are interbedded with laminated chalk (facies 3). units thicker than 1 m overlie facies 10 or facies 12–14. the lowermost unit is 8.15 m thick, overlies deformed bioturbated facies 12 and shows scattered burrows, although very few, throughout the unit. 2424 interpretation. the structureless appearance is considered a primary sedimentological feature, although it cannot be ruled out in all cases that the lack of sedimentary structures or burrows is simply due to a lack of colour contrast within the material. the facies is most commonly overlain by thinner laminated chalk units of facies 3, however, which may suggest a genetic relationship. also conspicuous is the common association with deformed or redeposited units. in addition, most units are burrowed at their top only, with burrows increasing in density toward the top, indicating that burrowing took place after deposition. the structureless appearance indicates rapid deposition excluding any sorting or structuring of the chalk ooze. the facies may thus be interpreted to have been deposited by mudflow, which is essentially a clast-free debris flow. the lack of clasts may be attributed to a lack of coherence in the source material, or that the flow was too dilute to possess sufficient shear strength to support clasts. the very thick unit in the lower part of the succession (plate 1, 3434.5–3426.4 m) is burrowed, although to a limited extent, throughout its 8.15 m of thickness, and is interpreted to have been deposited at a rate roughly equal to the rate of burrowing, which points towards deposition from a high-concentration suspension cloud or to the amalgamation of several flows. all observations taken together indicate that facies 9 beds were deposited by high-concentration gravity-driven suspension flows, which formed dilute visco-plastic flows during the final stages of deposition. facies 10: structureless matrixsupported chalk conglomerate description. this facies consists of a structureless chalk mudstone or more rarely wackestone matrix with dispersed fine sandto boulder-sized clasts, which are typically randomly distributed and oriented but may show coarse-tail grading (figs 19, 20). burrows are absent or restricted to the top of individual units. the facies makes up a total of 30.5 m distributed over 16 units, varying in thickness from 15 cm to 19 m. thin units generally occur within intervals dominated by facies 1–8. the clasts may be divided into chalk clasts, other lithic clasts, and bioclasts (fig. 21). chalk clasts are abundant, white to light grey, medium sandto bouldersized, although most are granuleto pebble-sized. they are mainly rounded to subangular. some chalk clasts show smear-structures ranging from minute flame-like structures extending subhorisontally from the clast surface into the matrix, over bidirectional double tails, to completely streaked-out clasts, forming thin white and commonly folded bands. incipient rip-up clasts, recognised as angular chalk clasts immediately above similarly shaped cavities in the underlying bed, have been observed at the base of redeposited units. many chalk clasts in thick units are considerably richer in fossils than the surrounding pelagic chalk, with bryozoan and small shell fragments making up the bulk of the fossils. a few hardground clasts have been observed. fossilrich chalk clasts may show surfaces with punctate irregularities reminiscent of borings, but these surface fig. 18. log example of structureless chalk (facies 9, see plate 1). although not bioturbated, a few isolated burrows are evident. 11261 ft 11262 11263 3432.5 3432.25 m 3433 burrows flint 2525 structures are moulds of small fossil fragments which were chipped off the clast during redeposition or possibly dissolved. other lithic clasts are rare and include green sandand siltstone clasts, shale chips, and small round to elongate glauconitised chalk clasts. bioclasts are dominated by bivalve shell debris, especially sand to fine pebble-sized inoceramid fragments. foraminifers make up an important but less conspicuous group, and bryozoan fragments occur in thicker units. other observed but rare bioclasts are from brachiopods, echinoderms, sponges, belemnites, and corals. a typical unit has a light grey, porous matrix, in contrast to the white and less porous medium sand to coarse pebble-sized chalk clasts. the matrix itself is structureless, but the clasts are coarse-tail graded in some units. one notable exception is the lowermost occurrence of the facies, which is a c. 5 m thick reddish coarse-tail graded inoceramid chalk wackestone with boulders composed of bioturbated marly chalk (facies 1 cm1 cm a b fig. 19. typical appearance of a structureless matrix-supported chalk conglomerate (facies 10; a: 3363.01–3362.89 m. b: 3247.59–3247.42 m (see plate 1). the sample in a was treated with light oil for colour-contrast enhancement. note the prevalence of subangular white chalk clasts. dark bioclasts (mostly inoceramid fragments) and a few large chalk wackestone clasts are also present (arrows). the crescent-shaped structure in the upper left part of a is a drill mark made from plug sampling. 2626 2), which most likely originated from the underlying lowermost cenomanian (fig. 22). the matrix is relatively fossil-rich and appears to be derived from a more shallow-water facies than the underlying and overlying chalk. the boulders are highly fractured and show numerous internal discordant surfaces. the fractures are not expressed by the boulder surface morphology, indicating that fracturing took place prior to emplacement. three 20–25 cm thick units in the 3411.6–3378.6 m (11 192 ft 9 in. – 11 084 ft 9 in.) interval, which is dominated by facies 1–8, constitute another variant of facies 10. they consist of coarse-tail graded, completely glauconitised medium sandto granule-sized, oval clasts in a structureless mudstone matrix. one of the units is directly overlain by facies 8, and the two others are overlain by facies 3. basal surfaces can only be examined in nine out of the sixteen units due to stylolites or poor core preservation; all but one of the nine units are sharply based. two of the basal surfaces are clearly erosional and two others are marked by thin shear zones. upper boundaries may be sharp or gradational, depending on the overlying facies. units with burrowed tops may grade into bioturbated chalk (facies 1) and units occurring within deformed intervals may grade into shear-deformed, matrix-supported chalk conglomerates (facies 13), although the transition may also be quite abrupt. upper boundaries are sharp when overlain by facies 5–8 or laminated chalk (facies 3). interpretation. the matrix-supported clasts indicate deposition by ‘freezing’ of a flow with some shear fig. 20. shear-banded chalk (facies 14, lower part) overlain (at c. 3250 m) by a structureless matrix-supported chalk conglomerate (facies 10); the two facies are separated by a stylolite surface. the colour banding in the upper part of the facies 14 unit shows tight to isoclinal folds. mud sand gr f m c pebbles 10656 10659 10662 10665 10668 10653 10655 10658 10661 10664 10667 10654 10657 10660 10663 10666 3247 m ft 3248 3249 3250 3251 chalk clasts and shell fragments shear banding 2727 strength, and the structureless matrix indicates total homogenisation during flow but sudden deposition in the absence of a sorting mechanism. the material must thus have possessed both fluid-like properties during flow and plastic properties during deposition. this type of visco-plastic, or bingham plastic, rheology is a wellfig. 21. clasts associated with facies 10, 11, 13 and 14. a: white subangular chalk clasts are the most common type of clast, but note also the large green marl clast in the lower part of the image and a greenish packstone clast in the uppermost part of the image. from 3422.65 m (see plate 1) b: common appearance of chalk clasts in facies 14. note the small size of the clasts and white streaks extending laterally from the clasts. the 3 mm long dark object is an inoceramid fragment. from 3325.25 m (see plate 1). c: exotic chalk clast rich in bioclasts (mostly small inoceramid fragments). from 3361.87 m (see plate 1). d: fragments of bryozoan colonies (arrows) within a structureless matrix-supported chalk conglomerate (facies 10). from 3247.42 m (see plate 1). b) a b c d 5 mm 5 mm 5 mm 5 mm 2828 documented feature of debris flows (johnson 1970), in which a relatively high-concentration sediment–water mixture, flowing under the influence of gravity, is able to support clasts much larger than the individual matrix components. chalk debris flows must have been formed by failure of chalk, whereupon relatively unconsolidated material was completely mixed with water until reaching a fluid state, whereas more consolidated and perhaps somewhat cemented material disaggregated to form chalk intraclasts. deeply buried, fully lithified chalk was thus rarely reworked into debris flows, and hardground clasts are rare exceptions. fig. 22. lower part of a structureless matrix-supported chalk conglomerate (facies 10) with boulder-sized clasts (dark tone) similar to the underlying bioturbated marly chalk (see also plate 1). the base of the unit just below 3440.5 m is marked by an argillaceous shear zone. the boulders were subjected to intense fracturing and welding prior to redeposition. mud sand gr f m c 11284 11285 11286 11287 11288 3440 3439.5 m ft 3440.5 fractures subvertical zoophycos boulder-sized clasts shell fragments burrows and faint lamination 2929 facies 11: clast-supported chalk conglomerate description. this facies is characterised by granule– pebble grade, clast-supported conglomerates but also includes pebbly grainstone beds. chalk clasts dominate, with subordinate shell debris (figs 23, 24). units belonging to this facies have sharp and loaded bases, whereas the upper boundary may be sharp or gradational, depending on the grain size. there is a range of variation within the ten units referred to facies 11. four units are 2–8 cm thick and are composed of roughly equal amounts of chalk clasts and shell debris in the coarse pebble to coarse sand fraction. these units show normal or inverse-to-normal grading. four other units are relatively fine-grained, with clast sizes ranging from granule chalk clasts to fine sand bioclasts with rare pebbles. the units are normally graded or inverseto-normally graded, with the rare pebbles occurring randomly without reference to the overall grading. they typically grade from basal granule to medium sand conglomerate or grainstone to medium–fine sand wackestone at the top. two of these units show distinct parallel stratification, whereas the two others are somewhat homogenised by burrowing. one of the units consists of well-defined rounded chalk clasts, whereas the others are mainly composed of highly compressed chalk clasts with diffuse boundaries, and numerous holes after dissolved grains. the chalk clasts commonly appear to be welded together and may be deformed by burrows or at grain–grain contacts. a spectacular example is provided by a 25 cm thick graded, coarse pebble to granule conglomerate bed composed entirely of chalk clasts (fig. 23). a texturally similar unit occurring between two facies 14 units shows heavily deformed boundaries. interpretation. the diagnostic character of facies 11 is the dominance of chalk clasts. it can be envisaged that significant numbers of relatively large chalk clasts may be produced by erosion and reworking of firm, buried chalk or of early diagenetic chalk nodules. liberation of the relatively large numbers of chalk clasts making up the conglomerates and grainstones of facies 11, however, would have required reworking of large volumes of chalk, and subsequent segregation of the mud and sand-to-pebble fractions. winnowing of ooze from nod ular chalk or from a chalk debrite could produce a lag conglomerate composed entirely of chalk clasts. the winnowing process would have involved removal 1 cm fig. 23. exceptional example of a graded, clast-supported chalk conglomerate (facies 11; 3269.21–3268.95 m, see plate 1). intergrain pore space is filled with cement and is largely devoid of mud. 3030 1 cm1 cm 1 cm a b c 3131 of enormous amounts of chalk ooze to produce conglomerates of the thicknesses observed in the mona1 core. it would take numerous cycles of debris flow or nodular hardening followed by complete removal of mud to produce thick layers of conglomerate or grainstone, because a thin residual lag would form an armoured layer sheltering the underlying chalk. furthermore, the lag would have to be reworked after its formation to produce the grading observed in facies 11. a more likely explanation is that the loss of mud occurred by extreme down-slope maturation of debris flows, whereby mud was expulsed from the flows (see krause & oldershaw 1979), until the point where the debris flows had essentially evolved into concentrated density flows as indicated by the overall normal grading of most units (fig. 25). grain flow dynamics affected the basal part of some flows during the final stages of deposition as indicated by the presence of basal inverse grading. facing page: fig. 24. clast-supported chalk conglomerates and pebbly grainstones (facies 11); all samples are largely composed of chalk clasts. a: normally graded and laminated grainstone mainly composed of coarse–fine sand-grade chalk clasts with some scattered chalk granules and pebbles (3343.93–3343.68 m, see plate 1). b: deformed conglomerate with clasts up to 5 cm across (3275.57– 3275.41 m, see plate 1). c: conglomerate (3361.79–3361.94 m, see plate 1); somewhat contorted due to compaction, but the overall grading is retained. the black colour is due to pyrite. facies 10facies 11 fig. 25. downstream maturation of debris flows whereby mud is progressively expelled from the flow until it consists of an upper mudflow and a basal turbidity/grain flow. inspired by krause & oldershaw (1979). 3232 facies 12: shear-deformed bioturbated chalk description. this facies may be regarded as a variant of facies 1, but it is described separately as it is important for the understanding of the mass-transport processes and the porosity variations. it is characterised by its generally biomottled but shear-deformed fabric (fig. 26). the biomottled fabric is evident but individual burrows are poorly preserved. planolites and chondrites can commonly be recognised by their morphology and configuration, but their margins show stepwise lateral off-sets on the submillimetre scale. well-defined shear zones, several centimetres thick, showing larger strain than the remaining unit have only been observed at the base of one unit. deformation was thus accommodated fig. 26. shear-deformed bioturbated chalk (facies 12; a: 3183.03–3182.85 m. b: 3159.99–3159.86 m, see plate 1). note the mottled fabric, which appears to have been slightly deformed. 1 cm 1 cm a b 3333 by limited displacement along numerous submillimetre-scale shear zones distributed rather evenly throughout the units. the amount of displacement along each of the small-scale shear zones may vary gradually within units. thus, undeformed bioturbated chalk (facies 1) may grade upwards into shear-deformed bioturbated chalk (facies 12), showing gradually increasing strain upwards to the point where the bioturbated fabric is hardly recognisable. units showing relatively large strain commonly contain small sand to granule grade intraclasts formed by the shearing. slightly more consolidated parts of the chalk apparently remained coherent during shearing of the surrounding chalk and hence produced undeformed but slightly displaced and sharp-bounded intraclasts. there thus appears to be a gradual transition between facies 12 and facies 14 with relatively few and small intraclasts, and these two facies are difficult to distinguish in the uppermost part of the core. there are nine individual occurrences of the facies in the core. the units are 25–660 cm thick and make up a total of 17.2 m. they are restricted to mass-transport fig. 27. photomicrographs of thin section of shear-deformed bioturbated chalk (facies 12; 3182.39 m, see plate 1) seen in plane polarised, transmitted light (a) and undercrossed polars (b). note the absence of intraparticle cement and thus preservation of porosity. a b 1 mm 1 mm 3434 intervals and are typically associated with facies 13 and 14. the boundaries to other facies can only be observed in a few cases due to stylolites and because the facies is mostly found in highly porous parts (fig. 27), where the core condition is relatively poor. however, it appears that the facies may have either sharp or gradational boundaries to other facies. interpretation. facies 12 represents small-scale plastic deformation of bioturbated chalk. the deformation may have occurred through several different processes. a common problem of core interpretation is to determine whether units occur in place between two slump units or they represent transported units within a single slump. the four upper facies 12 units almost certainly represent deformed bioturbated chalk within larger slumps, and are thus allochthonous, whereas the two lower facies 12 units were probably formed by deformation of bioturbated chalk by overriding debris flows, and therefore essentially represent in-place deformation (plate 1). one unit at 3274.2 m (10 743 ft) probably represents limited lateral creep of bioturbated chalk. two thin units, at 3218.4 m (10 559 ft) and 3182.7 m (10  442 ft), may be interpreted as bioturbated chalk deformed by overriding debris flows or slumps, or alternatively as intra-slump units. it is important to note that intraclasts may form through shear deformation of consolidated chalk. facies 13: shear-deformed matrixsupported chalk conglomerate description. this facies shares most of the general characteristics of facies 10, but differs in showing sheardeformation structures and in having a heterogeneous clast distribution (figs 28, 29). the deformation structures are seen as irregular colour bands in the matrix, which define open to tight subangular folds. several orders of smaller folds are superimposed upon larger folds and the boundaries between individual colour bands appear somewhat serrated and show small-scale flame structures. evidence for shear deformation on clast margins is relatively common and some clasts are sheared to the point at which they form thin white bands. clasts are randomly distributed on a decimetre scale, but are rather unevenly distributed on a larger scale, which may be due to larger-scale folding. there are 27 occurrences of the facies, which makes up a total of 50 m of the core. individual units are 20–960 cm thick and almost exclusively associated with mass-transport facies. most commonly, facies 13 shows gradual transitions to other shear-deformed facies (facies 12, 14 or facies 9 with deformation structures), and gradational transitions to facies 14 are particularly common (fig. 30). more rarely, the facies has rather sharp boundaries, which may be marked by a thin shear zone or bedding discordances. interpretation. this facies represents debrites deformed either by slumping or due to shearing by overriding flows or slumps. the precursor facies was similar to facies 10, but was plastically shear-deformed. the shear deformation most commonly occurred within larger-scale deformed units, involving several different facies. it is difficult to determine whether stacks of shear-deformed facies represent single or multiple deformation events, especially in a core. in the 3422.7– 3415.6 m (11 229 ft 4 in. – 11 206 ft) interval, the facies definitely represents three events as shown by burrowed surfaces within the succession. individual event beds were only a little more than 1 m to less than 4 m thick. in the lowermost bed, the degree of shear deformation increases upwards so that facies 10 grades into facies 13. in most cases, the deformation appears to have occurred subsequent to deposition, but it cannot be ruled out that some of the thin slumps composed of only facies 13, or facies 10 and 13 together, were deformed during the final stages of debris flow as the rheology of the material entered the plastic phase. in that case, the shear deformation was driven by the inertia of the debris flow itself, rather than by later additional gravityinduced stress. 3535 fig. 28. typical shear-deformed, matrix-supported chalk conglomerate (facies 13; 3197.44–3197.28 m, see plate 1). the facies appears similar to facies 10, but shows evidence for shear-deformation through relatively gentle folding. 1 cm 3636 fig. 29. typical shear-deformed, matrix-supported chalk conglomerate (facies 13, see plate 1) succession with evidence for both metreand centimetre-scale folding and a random clast distribution. mud sand gr f m c pebbles 11209 11210 11211 11212 11213 11214 3417 3416.5 m ft 3417.5 3418 burrows missing core chalk clasts and shell fragments colour banding 3737 fig. 30. succession showing a gradual upward transition from shear-deformed, matrix-supported chalk conglomerate to shearbanded chalk (facies 13–14). the transition at c. 3219 m is defined by a gradual, upward-increasing aspect ratio of folds with the boundary between facies 13 and 14 at 3219 m. the top part is bioturbated chalk (facies 1). mud sand gr f m c pebbles 10559 10560 10561 10562 10563 3219 3218.5 m ft 3219.5 biomottling chalk clasts colour banding shearbanding burrows 3838 1 cm fig. 31. shear-banded chalk (facies 14; 3344.41–3344.22 m, see plate 1) with relatively small, rounded chalk clasts. the colour banding shows the limbs of small isoclinal folds in the lower half of the sample. 3939 facies 14: shear-banded chalk description. this facies is characterised by the presence of matrix-supported chalk clasts, as in facies 10 and 13, but differs in showing an irregular colour banding (figs 31, 32, 33). the bands are 2–20 mm thick, and most are around 8 mm thick. individual bands are slightly wavy with somewhat irregular thickness. their boundaries have a skewed, serrated or flame-like appearance similar to those of facies 13. the bands commonly form thin recumbent isoclinal folds with closely spaced limbs, and the bands may also form larger, more open folds, especially when transitional to facies 13. the concentration of clasts varies greatly, both within and between individual units. they appear to show a random distribution within units and are mostly medium sandto granule-sized. larger clasts occur less commonly than in facies 10 and 13, and some units contain very few clasts. larger clasts commonly protrude through both base and top of single colour bands, but it is equally common to see colour bands that envelop clasts larger than the thickness of the band. some units contain distinct sandy layers up to a few centimetres thick, which have the same irregular appearance as the colour bands (fig. 33). it is difficult to distinguish units of facies 14 that have a low colour contrast and very few clasts from the most deformed units of facies 12. facies 14 is very common and makes up a total of 56 m distributed over 25 individual units, which are 15–705 cm thick. the facies is closely associated with facies 12 and 13. lower boundaries are either gradational to facies 13 or sharp and in some cases sheared into various other facies. upper boundaries are mostly gradational to slightly deformed facies 9 or facies 12 and 13, or sharp and in some cases discordant to facies 6, 9, 12, or 13. interpretation. this facies was formed through intense plastic shear deformation of chalk. the chalk must have been in an unlithified but somewhat dewatered and firm state, indicating a burial depth of about 10–200 m prior to deformation (mallon & swarbrick 2002). the close and gradational association with the folded facies 13 and the many examples of thin recumbent isoclinal folds indicate that the laminated appearance of the matrix is a type of shear banding. relatively gentle shear folding of firm but non-hardened chalk produced facies 13, whereas further shear deformation produced more folds and stretched already existing folds. this process caused a gradual horizontal lengthening and vertical shortening of the recumbent folds, and thus a higher aspect ratio. a core section through a succession of very thin recumbent folds with very large aspect ratios would display numerous beds or laminae, representing the limbs and only few fold hinges. it is quite possible that the process proceeded beyond the point of dismemberment at the fold hinges, thus further diminishing the likelihood of encountering a hinge in a core section. similar structures have been referred to as ‘shredded’ and have been attributed to ‘secondary shear-induced lamination’ in relation to slumping (kennedy 1980) and ‘flow-induced lamination’, which may form during the final stages of debris flows (nygaard et al. 1983). the lamination is clearly a product of plastic deformation and it is difficult to determine if the deformation within a debrite occurred during the final stages or after deposition of a fluidal debris flow. however, as with facies 13, it appears likely that debris flows entered a phase of plastic rheology during their final stages of flow, as also suggested in the interpretation by nygaard et al. (1983). the close association between facies 10, 13 and 14 indicates that facies 14 was typically produced by shear deformation of debrites (facies 10 and 13). however, facies 14 may also be produced by shear deformation of several other facies. units of facies 14 typically contain fewer and on average smaller chalk clasts than facies 10 and 13, mainly due to destruction of the clasts during intense shearing. it is also evident that chalk clasts may be produced directly by shear deformation of bioturbated chalk, as seen in facies 12, and a similar process may be expected to have occurred during the formation of facies 14. in that case, the ‘clasts’ are products of largely in-place brecciation of soft but firm chalk rather than true transported clasts. it is not possible to confidently distinguish the two in the present case, and the term clast is therefore maintained for simplicity. the distinction between facies 12 and 14 is often quite subtle, especially in the upper part of the core. in this interval, the similarity between relatively strongly sheared and clast-bearing facies 12 and 14 with weakly developed colour banding, suggest that the two facies are part of a continuum. in two other cases, at 3426.4 m (11 241 ft 7 in.) and 3218.1 m (10 558 ft 2 in.), facies 14 contains numerous deformed bioclastic sandy layers and appears to have been produced by deformation of a facies 1–8 dominated succession (fig. 33). 4040 1 cm a b 1 cm 1 cm c fig. 32. examples of shear-banded chalk (facies 14; a: 3257.02–3256.79 m. b: 3307.94–3307.84 m. c: 3325.22–3325.08 m, see plate 1). a shows relatively regular colour bands, whereas b and c show contorted, irregular colour bands. 4141 fig. 33. log example of shear-banded chalk (facies 14) containing deformed sandy layers, suggesting that it was formed by shear deformation of a pelagic/turbiditic succession (facies 1–8). mud sand gr f m c pebbles 11237 11238 11239 11240 11241 3426 3425.5 m ft brecciation chalk clasts shear-banding calcite veins 4242 slumps all the facies described above may have been incorporated into a slide or slump, yet none of the facies are diagnostic of slides and slumps. the recognition of slides and slumps must therefore be based on other parameters than facies alone. slides, defined as lateral transport of sediment along discrete shear planes with little or no internal deformation (e.g. stow et al. 1996), are difficult to recognise in cores. slumps are more easily recognised because they are internally deformed (stow et al. 1996), but it is often difficult to determine the number of depositional events represented by a succession of slumped facies. shear zones and undeformed burrow horizons may define the base and top of slumps, but even these criteria must be cautiously applied. apparently ‘normal’ burrowed/bioturbated chalk may represent undeformed but redeposited boulders, and shear zones may develop internally within slumps so that they do not define the base of a depositional unit. bypassing debris flows and slumps may cause deformation of the underlying strata, producing similar structures to those formed in a slump. this type of deformation did not involve subaqueous motion of the deformed sediment and should be distinguished from true slides and slumps. fourteen intervals (numbered s1–s14 in ascending stratigraphic order) are interpreted to represent slumps (table 2). several of these contain thin zones of undeformed bioturbated chalk. in s14, the undeformed zones appear to be overturned, judging from the general alignment of the biomottled fabric and the bifurcation direction of chondrites. in s12, a slightly deformed bioturbated zone is sharply based, which would not be the case if it represented the bioturbated top of a slump. overturned and sharply based bioturbated zones represent undeformed boulders or plugs within slumps rather than the burrowed top of a slump. in other cases, the zones appear to be correctly oriented and the boundaries are not preserved or marked by stylolites. such zones may represent pelagic deposition between slump events or undeformed boulders. s6, s7, s10 may represent multiple events, and there are thus 14 –18 slump events represented in the succession, and individual slump units are 2 m to more than 42 m thick. debrites appear to be the most common precursor facies involved in slumping and there are several possible explanations for this. the slump structures may have formed during the transition between fluid and plastic behaviour as an integral part of debris flow deposition. slumps and debris flows probably developed due to similar preconditions and triggers, meaning that they occur in the same temporal and spatial setting. debrites lacked the efficient dewatering of more slowly deposited and bioturbated sediment, and were therefore more susceptible to renewed slope failure. 4343 table 2. slump units: distribution and characteristics depth thickness facies precursor facies remarks s1 3426.56–3424.17 m 2.39 m 14 pelagite/ highly shear-deformed chalk with few and small (11242 ft–1234 ft 2 in.) turbidite succession chalk clasts and deformed sandy packstone layers. burrowed at top. s2 3422.70–3420.52 m 2.18 m 13 debrite debrite shear-deformed on the sea floor (slump) (11229 ft 4 in.–11222 ft 2 in.) 10 or in the subsurface by subsequent debris flow. capped by turbidite. s3 3420.52–3415.39 m 5.13 m 13 debrite basal shear surface and burrows at top. (11222 ft 2 in.–11205 ft 4 in.) succeeded by four amalgamated mudflow units (facies 9). s4 3398.14–3390.52 m 7.62 m 14 pelagite? highly shear-deformed slump with few and (11148 ft 9 in.–11123 ft 9 in.) 13 relatively small chalk clasts, suggesting a non debrite origin. internal zone of less pronounced deformation (facies 13). s5 3382.90–3378.63 m 4.27 m 14 debrite basal shear zone. (11098 ft 9 in.–11084 ft 9 in.) 13 s6 3323.52–3317.14 m 8.38 m 13 lower half one or two slump events. possibly one large (10910 ft 6 in.–10883 ft) 12 probably pelagite, slump with internal zone/boulders of 14 upper part debrite bioturbated chalk, or three discrete slumps separated by pelagic deposits. s7 3314.85–3307.69 m 7.16 m 14 debrite facies 1 unit appears to be inverted (10875 ft 6 in.–10852 ft) 13 suggesting it is a large clast within a 1 large slump. if not, the unit separates 10 two discrete flow units. s8 3307.08–3302.81 m 4.27 m 13 pelagite basal part of unit was probably (10850 ft–10836 ft) 14 originally a debrite. capped by turbidite. s9 3282.39–3275.69 m 6.70 m 13 debrite/ basal shear zone. capped by (10769 ft–10747 ft) 14 pelagite turbidite. 9 succession s10 3275.46–3272.64 m 2.82 m 1 pelagite bioturbated interval (facies 1) may (10746 ft 3 in.–10737 ft) 13 represent pelagic deposition between 14 two slump events, or an undeformed 3 plug or boulder within a single slump. s11 3257.09–3250.01 m 7.08 m 14 debrite (10686 ft–10662 ft 9 in.) s12 3227.43–3216.81 m 10.62 m 14 debrite/ highly shear-deformed succession. originally basal (10588 ft 8 in.–10553 ft 10 in.) 13 pelagite/ debrite and upper pelagic/turbidite interval. facies 12 turbidite 12 interval shows sharp lower boundary, indicating 10 succession that it represents a boulder within the slump. s13 3202-23–3188.00 m 14.23 m 13 debrite (10506 ft–10459 ft 4 in.) 14 s14 3184.70–3142.49 m (min.) >42.21 m 12 pelagite/ bioturbated chalk intervals are (10448 ft 6 in.–10310 ft) 14 debrite overturned. several internal 13 succession discordant surfaces. 1 10 4444 genetically related facies facies 1–4 (table 1) are all interpreted to be the product of pelagic deposition, possibly interrupted by deposition of thin event beds, and are collectively termed pelagites. facies 5–8 and facies 11 are all interpreted as having been deposited by turbidity flows and are collectively termed turbidites. facies 7 and 8 are interpreted as equivalent to the basal part of facies 6, and facies 5 is equivalent to the uppermost part of facies 6. facies 7 and 8 thus, in theory, represent the most high-energy (or proximal) turbidites and facies 5 the most low-energy (or distal) turbidites (fig. 34). facies 2 (laminated chalk) may be incorporated into this scheme as representing the most distal or low-energy facies related to turbidity flows. facies 11 was produced by down-slope evolution of debris flows (fig. 25). limited shear deformation produced distinctly deformed facies from which the precursor facies can still be deduced (facies 12 and 13), whereas intense or continued shear deformation produced the shear-banded facies 14 from which the precursor facies cannot be deduced in all cases. facies 9, 10 and all beds occurring within the 14 slump units were deposited by mass transport as defined by moscardelli & wood (2008) and are collectively termed mass-transport deposits. 4545 facies 5 facies 6 facies 8facies 7 facies 8facies 7 facies 5 facies 6 a b fig. 34. hypothetical relationship between turbidite facies 5–8, interpreted to represent distance to source area (a) or variable basin gradients (b). 4646 facies associations and stratigraphic development facies associations three facies associations are defined on the basis of a dominant or characteristic group of facies. they comprise intervals dominated by pelagic deposits, mixed pelagite/turbidite intervals, and intervals dominated by mass-transport deposits. facies association 1 (pelagic chalk) description. facies association 1 (fa 1) occurs in six intervals and makes up 26% of the upper cretaceous succession (fig. 35, plate 1). it is dominated by pelagic facies, mainly facies 1 with subordinate facies 2–4. intrapelagic erosion surfaces are common in the lower part of the core and turbidites are rare and occur dispersed through the pelagic intervals. mass-transport facies are very rare, and there appears to be no systematic ordering of facies within intervals belonging to this facies association. interpretation. fa 1 represents prolonged periods of relatively stable pelagic deposition, only rarely interrupted by redepositional or erosional events caused by bottom currents or benthic storms. although the association accumulated mainly under well-oxygenated benthic conditions, dysoxic bottom conditions occurred periodically. the rarity of redeposited facies indicates tectonic quiescence and a low-gradient basin topography. facies association 2 (turbiditic chalk) description. facies association 2 (fa 2) occurs in two intervals and makes up 23% of the upper cretaceous succession (fig. 35, plate 1). it is represented by common to abundant turbidites, alternating with pelagites and rare to common mass-transport facies. bioturbated chalk (facies 1) is the most common pelagite, but laminated chalk (facies 3) dominates in the lowermost part of the lower interval and occurs locally throughout both intervals. intrapelagic erosion surfaces are common in the upper part of the upper interval. turbidites show an overall thinning-upward development, but shorter subintervals may show either thinningor thickeningupward trends. the dominant turbidite facies changes upwards from facies 6 to facies 7, to facies 8 and finally to facies 5 (fig. 36). interpretation. fa 2 represents a relatively dynamic depositional environment with pelagic deposition interrupted by recurring slumps, turbidity, mud and debris flows. benthic dysoxia occurred episodically, but most of the facies association accumulated under well-oxygenated benthic conditions. the association is interpreted to represent deposition in a basinal setting. facies association 3 (mass-transport chalk) description. facies association 3 (fa 3) occurs in six intervals and makes up 51% of the upper cretaceous succession (fig. 35, plate 1). it is overwhelmingly dominated by mass-transport facies with only a few occurrences of non-reworked pelagites and turbidites. deformed facies and debrites constitute most of fa 3, and most of the facies occur in inferred slumps. a systematic ordering of facies is not observed within intervals belonging to fa 3. interpretation. fa 3 represents debris flow, mudflow and intervening pelagic deposition, and all of the resultant facies were commonly reworked into slumps. the association is interpreted to represent relatively unstable periods of the depositional history. facing page: fig. 35. simplified log of the studied succession showing the distribution of facies associations and the subdivision into stratigraphic intervals a–e discussed in the text. alb.: albian. cen.: cenomanian. camp.: campanian. 4747 up pe r al b. c en . * * middle coniacian – lower santonian lo w er m aa st ric ht ian up pe r c am p. up pe r m aa st ric ht ian facies association description interpretation 3184.70 m (10448 ft 6 in.) 3191.05 m (10469 ft 4 in.) 3202.23 m (10506 ft) 3216.81 m (10553 ft 10 in.) 3257.09 m (10686 ft) 3272.64 m (10737 ft) 3282.39 m (10769 ft) 3290.77 m (10796 ft 6 in.) 3301.70 m (10832 ft 4 in.) 3324.10 m (10905 ft 10 in.) 3352.11 m (10997 ft 9 in.) 3411.55 m (11192 ft 9 in.) 3440.58 m (11288 ft) fa 3 fa 3 fa 3 fa 3 fa 3 fa 3 fa 2 fa 2 fa 1 fa 1 fa 1 fa 1 fa 1 fa 1 upward transition from laminated marl to bioturbated marly chalk. dominated by debrites, mud flow deposits, and deformed facies. also minor pelagites and turbidites. pelagites alternating with common turbidites and some mass-transport deposits. turbidites show an overall thinning-upward trend and an upward change from dominantly high-energy to lowenergy turbidite facies; individual intervals bounded by mass-transport facies commonly show opposite trends. pelagites with some intrapelagic erosion surfaces, turbidites and thin debrites. dominated by mass-transport deposits alternating with pelagites. turbidites are very rare and deformed facies become increasingly dominant and thicker upwards. pelagic deposition interrupted by increasingly large and frequent mass transport events. relatively quiet deposition in a low-gradient basin. deposition of prograding and aggrading intervals during gradually decreasing basin gradients. at least eight discrete mass transport events, probably due to uplift of the mona ridge. hemipelagic/pelagic deposition. increasing oxygenation and reduced influx of siliciclastic material. e d c b a mass-transport chalk pelagic chalk turbiditic chalk 11300 11270 11240 11210 11180 11150 11120 11090 11060 11030 11000 10970 10940 10910 10880 10850 10820 10790 10760 10730 10700 10670 10640 10610 10580 10550 10520 10490 10460 10430 10400 10370 103403150 m ft 3160 3170 3180 3190 3200 3210 3220 3230 3240 3250 3260 3270 3280 3290 3300 3310 3320 3330 3340 3350 3360 3370 3380 3390 3400 3410 3420 3430 3440 depth 4848 stratigraphic development the succession belongs to the upper cretaceous with the exception of the lowermost 1.8 m (bailey et al. 1999) (fig. 35, table 3). the cored section extends upwards to 9.73 m below the cretaceous–tertiary boundary, which is reported to occur at 3132.7 m (10 278  ft) (bailey et al. 1999). the biostratigraphical zonation defined by bailey et al. (1999) was calibrated to the chronostratigraphic scheme of ogg et al. (2004). more a b d ec c occurrence unit thickness (cm) facies 5 facies 6 facies 7 facies 8 turbidites ipes 0 20 40 10300 10400 10500 10600 10700 10800 10900 11000 11100 11200 11300 3150 m ft 3200 3250 3300 3350 3400 a a b d ec 0 10 20 30 40 50 60 70 80 90 100 110 120 130 c um ul at ive th ick ne ss (m ) pelagic turbidites mudflow deposits debrites slumps 10400 ft10600108001100011200 3150 m32003250330033503400 10400 ft10600108001100011200 3150 m32003250330033503400 b c um ul at ive th ick ne ss (m ) 0 1 2 3 4 5 6 7 8 9 10 turbidites facies 5 facies 6 facies 7 facies 8 e d c b a fig. 36. a: cumulative thickness of pelagites, turbidites, mudflow deposits, debrites and slumps. note the upward change from mixed deposits to pelagite-dominated deposits and to slump-dominated deposits with intervening pelagites. debrites are under-represented in this plot as they were typically incorporated into slumps. b: cumulative thickness of turbidites. note the upward change in the dominant facies from facies 6 through 7 and 8 to 5. this suggests a prograding to aggrading basin development, but note that according to bailey et al. (1999) this interval spans a considerable hiatus (compare with figs 34, 37). c: thickness of turbidites (left panel) and occurrences of individual turbidite facies and intrapelagic erosion surfaces (ipes) (right panel). the plot illustrates the overall upward decreasing importance of turbidites, and the right panel shows that the upward change in dominant facies illustrated in b is succeeded by an increase in the occurrence of intrapelagic erosion surfaces. stratigraphic intervals a–e from fig. 35 4949 than half of the time represented by the studied succession is represented by hiatuses, and little more than 10% of the time is represented by pelagic deposits (fig. 37). the pelagic intervals show common signs of erosion and the remaining succession is composed of event deposits, which can be considered to have formed geologically instantaneously. this implies that 90% of the late cretaceous epoch is not represented in the succession. there is, however, great uncertainty involved in applying biostratigraphical markers within a succession that is strongly influenced by redeposition. based on the available data, however, the succession is subdivided into five intervals (a–e on figs 35–37, 39): (a) the lower 5.5 m of the succession (3446.1–3440.6 m, 11 306 ft – 11 288 ft) belongs to fa 1 and is of latest albian – cenomanian age (fig. 35). the interval shows an upward transition from laminated marl (facies 4) to bioturbated marly chalk (facies 2), which is interpreted to represent increased benthic oxygenation, enhanced production of pelagic carbonate and reduced siliciclastic influx. (b) this is succeeded by a 29 m thick interval (3440.6–3411.6 m, 11 288 ft – 11 192 ft 9 in.) belonging to fa 3 and referred to the cenomanian – late campanian age range. it represents at least eight individual mass-transport events, including a debrite with boulders of the older marly chalk in a matrix derived from relatively shallow-water carbonates, as well as two turbidites and minor intervening pelagites. this interval is interpreted to represent significant tectonic activity and uplift of the mona ridge. non-deposition or erosion thus characterised the cenomanian – late campanian interval, only interrupted by mass-transport events (fig. 38a, b). the overall accumulation rate for the period is 0.1 cm/ka, never exceeding 0.3 cm/ ka, with pelagic accumulation being an order of magnitude lower. for comparison, chalk group accumulation rates average around 2–2.5 cm/ka and may reach 12 cm/ka (ehrman 1986; schönfeld et al. 1996; surlyk et al. 2003). the cenomanian – late campanian thus represents 24 ma characterised by sediment bypass or erosion, with deposition being almost exclusively by mass transport (fig. 38c). (c) the overlying 59.4 m thick interval (3411.6– 3352.1 m, 11 192 ft 9 in. – 10 997 ft 9 in.) of late campanian – early maastrichtian age is a complex succession belonging to fa 2 in which nearly all facies recognised are represented. pelagic facies table 3. chronostratigraphic subdivision* chronostratigraphy upper maastrichtian (uc20)† upper maastrichtian (uc19) depth age (ma) 3132.73 m (10278 ft) 3288.79 m (10790 ft) 3332.23 m (10932 ft 6 in.) 65.5 68.6 69.2 lower maastrichtian (uc18) 3386.56 m (11110 ft 9 in.) 70.4 upper campanian – lower maastrichtian (uc17) 3416.20 m (11110 ft 9 in.) 75.2 upper campanian (uc16) 3386.56 m (11208 ft) 75.8 lower santonian – upper campanian (uc12–15) 3416.20 m (11208 ft) 84.5 middle coniacian – lower santonian (uc10–11) 3426.56 m (11242 ft) 88.0 turonian – middle coniacian 3426.56 m (11242 ft) 93.5 cenomanian 3444.09 m (11299 ft 6 in.) 99.6 *subdivision based on the biostratigraphic analysis of bailey et al. (1999) with the addition of chronometric ages from ogg et al. (2004) †uc = upper cretaceous nannofossil zones from burnett (1998) 5050 are dominated by bioturbated chalk (facies 1), but laminated chalk (facies 3) occurs sporadically throughout and dominates the lowermost part, which may be interpreted as a shallowing-upward trend. turbidites are common to abundant and show a general thinning-upward trend (fig. 36), which is interpreted to represent gradually decreasing sea-floor topography during progressive basin filling. it should be noted that this interval encompasses a nearly five million-year hiatus according to the available biostratigraphy. thus, it cannot be determined if the lithological interval represents completely separate depositional systems or a long-term depositional development interrupted by significant erosion. on a smaller scale, subintervals showing either thickeningor thinningupward trends represent prograding or aggrading trends superimposed on the general development. the terms prograding and aggrading are used here to describe the architectural development of the basin-fill, with no implications concerning the position of basin margins or shorelines. there is an upward change in the dominance of the different turbidite facies in the order: facies 6, 7, 8, and 5, which is interpreted to represent an increase in turbidity flow energy in the lower part of the interval and a gradual decrease in turbidity flow energy in the remainder of the interval (fig. 34). the stratigraphical development of turbidite facies suggests an initial rapid basinward shift of facies due to prograding basin infill, which represents a short period with a relatively high sedimentation rate (4.9 cm/ka) in the late campanian. this was followed by five million years with non-deposition or erosion and subsequent gradual shift towards more low-energy turbidite facies and to aggrading basin infill. intrapelagic erosion surfaces become common in the upper part of the interval. (d) the overlying 28 m thick interval (3352.1–3324.1 m, 10 997 ft 9 in. – 10 905 ft 10 in.) of early – late maastrichtian age belongs to fa 1 and is dominated by bioturbated chalk with some intrapelagic erosion surfaces, relatively rare turbidites and thin debrites. this interval is interpreted to represent relatively quiescent tectonic conditions in a lowgradient basin setting. (e) the remaining 181 m of the studied succession essentially comprises bioturbated chalk and massfig. 37. simplified log of the studied succession plotted against time. more than half of the late cretaceous time interval is represented by hiatuses, and only 10% is represented by pelagites. 95 90 85 80 75 70 c en om an ian tu ro ni an c on iac ian sa nt on ian c am pa ni an m aa st ric ht ian ag e (m a) e d c b a mass-transport chalk pelagic chalk turbiditic chalk hiatus 5151 transport facies of late maastrichtian age, intervals referred to fa 1 and fa 3. the only exception is a 10.9 m thick interval (3301.7–3290.8 m, 10 832 ft 4 in. – 10 796 ft 6 in.), which belongs to fa 2 and consists of bioturbated chalk, mudflow deposits, thin laminated chalk units, some turbidites and debrites. by far the greater part of the succession accumulated in this period (fig. 39), when bulk accumulation rates were as high as 7.2 cm/ka, and pelagic accumulation rates were up to 3.6 cm/ka. the late maastrichtian was thus characterised by relatively quiet pelagic deposition interrupted by large-scale redeposition events. the upward thickening and increasing abundance and dominance of slumps, and probably also debrites, may be due to a progressively shorter distance to the source and deposition further up the flank of the basin margins, or to a gradual increase in the frequency and intensity of triggering events. the stratigraphic development may thus be summarised as follows: (a) a latest albian – cenomanian gradual shift from dominantly hemipelagic, siliciclastic to dominantly pelagic carbonate deposition; (b) a long cenomanian – late campanian period associated with tectonic activity and mass transport; (c) late campanian – early maastrichtian basin filling dominated by pelagic and turbidite deposition; (d) early – late maastrichtian basin filling dominated by pelagic deposition; (e) late maastrichtian basin filling interrupted by increasingly frequent and voluminous mass-transport deposition. 0.0 50 100 150 200 250 300 95 90 85 80 75 70 95 90 85 80 75 70 95 90 85 80 75 70 age (ma) age (ma) age (ma) 0.0 0.3 0.3 0.0 4.9 4.5 7.2 5.0 0.0 cen. tur. con. sa. camp. maastr. cen. tur. con. sa. camp. maastr. cen. tur. con. sa. camp. maastr. 0 20 40 60 80 100 pe lag ic cu m ul at ive th ick ne ss (m ) 0.0 0.03 0.06 0.0 0.9 0.0 3.6 2.4 1.2 0 20 40 60 80 100 120 pelagic turbidites slumps debrites mudflow deposits a b c c um ul at ive th ick ne ss (m ) c um ul at ive th ick ne ss (m ) fig. 38. cumulative thickness plotted against time for the whole succession (a), pelagites (b) and genetic units (c). numbers on the curves are accumulation rates in cm/ka. the plots illustrate how most of the succession accumulated during the maastrichtian, whereas little is preserved from deeper intervals of the upper cretaceous. the low accumulation rates during the cenomanian– campanian probably reflect the presence of numerous hiatuses below chronostratigraphic resolution. cen.: cenomanian. tur.: turonian. con.: coniacian. sa.: santonian. camp.: campanian. maastr.: maastrichtian. 5252 thickness (m) 0 50 100 150 c en om an ian tu ro ni an c on iac ian sa nt on ian c am pa ni an m aa st ric ht ian a b c d e mass-transport chalk pelagic chalk turbiditic chalk 95 90 85 80 75 70 ag e (m a) fig. 39. simplified log of the studied succession plotted against time and thickness of defined stratigraphic intervals, illustrating the prevalence of maastrichtian and poor preservation of deeper upper cretaceous levels. 5353 discussion basin development and depositional history the earliest cenomanian was characterised by the transition from marl to chalk deposition, interrupted by mass-transport events, probably caused by uplift of the mona ridge (figs 40a, 41a) associated with the basin inversion which shifted the local depocentre eastwards from the jeppe basin to the karl basin (vejbæk & andersen 2002). less than 30 m of sediment, largely mass-transport facies, are preserved in the cenomanian – upper campanian succession, indicating that the mona-1 area was subjected to continuous or pulsed removal of sediments due to uplift of the mona high and subsidence of the karl basin in which deposition was focussed. the karl basin was filled during the late campanian and deposition continued across the karl basin, the mona high area and the jeppe basin. during this period, the mona-1 area was subject to rapid accumulation of pelagic chalk, thin mass-transport chalk, and interbedded turbidites in an overall prograding pattern (figs 40b, 41b). the absence of the upper campanian – lower maastricht ian nannofossil biozone uc17 may indicate erosion due to an inversion pulse (vejbæk & andersen 2002) or a significant sea-level drop (fig. 42) near the campaw e w e w emona-1 mona-1 mona-1 c. 1 km maastrichtian lower cretaceous cenomanian–campanian pre-cretaceous karl basinjeppe basin c. 100 m c b a fig. 40. interpreted basin evolution based on sedimentological interpretation and seismic section (see fig. 3). white dashed lines illustrate general reflection patterns, black dashed line represents approximate base of the prograding upper campanian succession. a: basin configuration in the latest early cretaceous, showing the lower cretaceous to be largely restricted to the jeppe basin. b: basin configuration in the latest campanian resulting from cenomanian–campanian relative uplift of the mona ridge and subsidence of the karl basin. the dashed black line represents the approximate onset of prograding pelagite/turbidite deposition across the mona ridge. older deposits form a thick succession in the karl basin, but are only represented by 30 m of mass-transport deposits and minor pelagites on the mona ridge. c: basin configuration in the latest maastrichtian. the maastrichtian succession consists of a lower aggrading pelagite/turbidite part and an upper part which is dominated by mass-transport deposits. 5454 nian–maastrichtian boundary. an aggrading pelagic/ turbidite succession was deposited at the margin of the karl basin at mona-1 during the early maastrichtian, followed by pelagic and debris flow deposition with recurrent large-scale slumping during the late maastricht ian, probably due to inversion and uplift of the highs east of the karl basin (figs 40c, 41c). it is tempting to link the depositional history to tectonic phases or sea-level variations. both have probably exerted a major influence on the depositional processes and patterns, but it should be noted that the studied succession is characterised by hiatuses and redeposited units so that almost any postulated sea-level or tectonic event could be roughly correlated to a previously reported event. there are few other wells available in a b c fig. 41. interpretation of basin development during the late cretaceous in the mona-1 area. a: cenomanian– campanian. the mona-1 area was subjected to continuous or pulsed erosion due to uplift of the mona ridge or subsidence in the karl basin. only 30 m of sediment, mainly mass-transport deposits are preserved in mona-1. the transport direction is speculated to be eastward because mona-1 is situated slightly eastward, down-slope from the culmination of the mona ridge. b: campanian – early maastrichtian. the karl basin was filled and the mona-1 area was subject to rapid accumulation of a prograding sequence of turbidites and pelagites, perhaps interrupted by an inversion pulse at the campanian–maastrichtian boundary. c: late maastrichtian. inversion and uplift of the highs east of the karl basin caused the depositional environment in the mona-1 area to be dominated by debris flows and recurrent large-scale slumping. 5555 kominz et al. (2008). best estimate (solid), smoothed (broken). haq et al. (1987). short term (solid), long term (broken) sea level 95 90 85 80 75 70 c en om an ian tu ro ni an c on iac ian sa nt on ian c am pa ni an m aa st ric ht ian ag e (m a) risefall mass-transport chalk pelagic chalk turbiditic chalk hiatus fig. 42. simplified log of the studied succession plotted with the eustatic sea-level curves of haq et al. (1987) and kominz et al. (2008). the curve of kominz et al. (2008) has been simplified for clarity and does not show the margin of error or the origin of the data as in the original version. 5656 the study area, and no drill cores covering the chalk group. attempts at linking chalk facies to well-log patterns have yielded rather unconvincing results, which inhibit the correlation of depositional patterns to other wells. thus, further data and studies are required to place the depositional history of the mona-1 succession within a regional framework on other than the most general level as outlined above. stacking pattern and predictability nygaard et al. (1983) proposed a model for the spatial and/or temporal evolution of chalk redeposition pro cesses, inspired by middleton & hampton (1976). according to the model, flow transformation causes gravity flows to evolve from slide/slump via debris flow and mudflow to turbidity flow. the vertical order of redeposited facies in chalk successions is generally considered chaotic and unpredictable (surlyk et al. 2003), which may have several causes: (1) the redeposition processes are line sourced, (2) the redeposition processes may be initiated at any location on the slope, (3) the dimensions of mass-transport flows are highly variable ranging from small and local to enormous regional-scale events, (4) down-slope flow transformation is highly variable and non-systematic, and (5) different flow conditions may co-exist within a single flow, so that one part of the resulting deposit will be recognised as a slump deposit, whereas another part will be characterised as a debrite. even if all these causes are valid, there should be a certain probability of encountering particular facies on a particular part of the slope; slumps should thus be more common in the up-slope direction. the chalk group in mona-1 shows a well-defined stacking pattern, although punctuated by numerous hiatuses, comprising a lower basinal succession characterised by a prograding and aggrading turbidite development, overlain by a slope succession dominated by thick slumps and debrites. if similar patterns exist in other chalk settings they may not have been recognised for a number of reasons. mona-1 is an abnormally long core spanning almost the entire chalk group, whereas most cores are much shorter and perhaps cover too little of the succession to detect a subtle upward trend in facies evolution. the turbidites defining the prograding/aggrading pattern constitute volumetrically minor parts of the succession, and the debrites and slumps were most probably line-sourced. thus, neither turbidites nor mass-transport facies produced geomorphological and bathymetrical elements that were sufficiently dramatic to allow recognition in seismic profiles. knowledge about the spatial geometry of the masstransport complexes is imperative for the further development of a predictive tool. well logs, principally sonic logs, have been utilised to delineate the distribution of mass-transport complexes in the north sea chalk (hatton 1986). the apparent relationship between masstransport facies and high porosity was extended to the assumption that high-porosity intervals, as revealed by high sonic log values (low velocity), could be taken to represent mass-transport facies, thus involving a certain degree of circular reasoning. this implies that the depositional process is the only or dominant influence on porosity, which is clearly an invalid assumption. almost all correlations of hatton (1986) were apparently made between wells with no core control of facies. it was further assumed that high-porosity intervals in wells more than 20 km apart were laterally connected. in other words, if the high-porosity zones did in fact represent mass-transport facies, it was apparently further assumed that they represented the same sedimentary body. such conclusions thus have far-reaching significance for the understanding of the depositional system, but they may be based on invalid assumptions. further refinement of geophysical methods and greater integration with sedimentological data are necessary in order to constrain the spatial geometries and connectivity of redeposited chalk bodies. 5757 conclusions this analysis of chalk in the exceptionally well-cored mona-1 borehole has resulted both in the detailed documentation and interpretation of a wide spectrum of facies types and in the elucidation of the depositional evolution of the mona ridge area in the late cretaceous. (a) the 303 m thick chalk upper cretaceous succession in the mona-1 core comprises 14 facies representing pelagic deposition, slumping, debris flow, mudflow and turbidity flow. (b) pelagic sedimentation probably took place by pelagic fall-out alternating with deposition of thin event beds, and the deposits were thoroughly bioturbated if deposited during periods of full oxygenation of the benthic realm. (c) four turbidite facies are recognised and interpreted as having formed by resuspension of unconsolidated chalk ooze on a slope; these different turbidite facies may be related in a high-to-low energy framework. (d) clast-supported chalk conglomerates are interpreted to have been directly related to the down-slope evolution of debris flows, which are represented by matrix-supported conglomerates. (e) most structureless chalk units were deposited by high-concentration, gravity-driven mudflows, but some burrowed structureless units may have been deposited by rapid fall-out from high-concentration suspension flows. (f) limited shear deformation produced distinct quasi-facies from which the precursor facies can be deduced, whereas intense or continued shear deformation produced a shear-banded quasi-facies from which the precursor facies cannot be interpreted in all cases. (g) more than 40% of the succession was affected by slumping; 14 –18 slump events are recognised, and debrites appear to be the most common precursor facies involved in slumping. (h) the stratigraphic evolution of facies records: (1) a gradual shift from dominantly hemipelagic siliciclastic to dominantly pelagic carbonate deposition in the earliest cenomanian, interrupted by mass transport and probably uplift of the mona ridge; (2) continuous or pulsed erosion and associated mass transport, probably due to uplift of the mona ridge, and filling of the karl basin during the cenomanian – late campanian; (3) pelagic and prograding turbidite basin filling in the late campanian; (4) probable uplift and erosion of the mona ridge close to the campanian–maastrichtian boundary; (5) aggrading basin filling dominated by pelagic deposition interrupted by increasingly frequent and voluminous mass-transport deposition during the maastrichtian. acknowledgements the project was funded by the danish council for independent research – natural sciences (fnu), maersk oil, and the faculty of science, university of copenhagen. thanks are extended to maersk oil staff henriette agersnap, dann graulund, ulla hoffman and henrik tirsgaard for assistance during the project. martin skov andreasen, lars ole boldreel, liesbeth breesch, janus bruun christiansen, nicolas thibault (all university of copenhagen), ida fabricius (dtu), claus andersen (geus) and emma sheldon (geus) are thanked for advice and discussion. finn jakobsen (geus) is especially thanked for providing seismic data and interpretation. we also thank haydon w. bailey and liam gallagher (both network stratigraphic consulting ltd) for comparing the biostratigraphical zonation with an updated database. the useful comments and suggestions from the referees, maurice e. tucker and frans van buchem, are gratefully acknowledged. 5858 references anderskouv, k., damholt, t. & surlyk, f. 2007: late maastrichtian chalk mounds, stevns klint, denmark: combined physical and biogenic structures. sedimentary geology 200, 57–72, http:// dx.doi.org/10.1016/j.sedgeo.2007.03.005 bailey, h. et al. 1999: a joint chalk stratigraphic framework. joint chalk research phase v, 203 pp. stavanger: norwegian petroleum directorate. bambach, r.k. 1993: seafood through time: changes in biomass, energetics, and productivity in the marine ecosystem. paleobiology 19, 372–397. bouma, a.h. 1962: sedimentology of some flysch deposits: a graphic approach to facies interpretation, 168 pp. amsterdam: elsevier. bramwell, n.p., caillet, g., meciani, l., judge, n., green, m. & adam, p. 1999: chalk exploration, the search for a subtle trap. in: fleet, a.j. & boldy, s.a.r. 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(eds): the millennium atlas: petroleum geology of the central and northern north sea, 45–59. london: geological society. ziegler, p.a. 1990: geological atlas of western and central europe, 2nd edition, 239 pp. the hague: shell international petroleum maatschappij b.v. http://dx.doi.org/10.1016/0037-0738(80)90052-4 http://dx.doi.org/10.1016/0037-0738(80)90052-4 http://dx.doi.org/10.1306/07100808035 6161 6262 de nationale geologiske undersøgelser for danmark og grønland (geus) geological survey of denmark and greenland øster voldgade 10, dk-1350 copenhagen k, denmark the series geological survey of denmark and greenland bulletin started in 2003 and replaced the two former bulletin series of the survey, viz. geology of greenland survey bulletin and geology of denmark survey bulletin. some of the twenty-one volumes published since 1997 in those two series are listed on the facing page. the present series, together with geological survey of denmark and greenland map series, now form the peer-reviewed scientific series of the survey. geological survey of denmark and greenland bulletin 1 the jurassic of denmark and greenland, 948 pp. (28 articles), 2003. edited by j.r. ineson & f. surlyk. 500.00 2 fish otoliths from the paleocene of denmark, 94 pp., 2003. by w. schwarzhans. 100.00 3 late quaternary environmental changes recorded in the danish marine molluscan faunas, 268 pp., 2004. by k.s. petersen. 200.00 4 review of survey activities 2003, 100 pp. (24 articles), 2004. edited by m. sønderholm & a.k. higgins. 180.00 5 the jurassic of north-east greenland, 112 pp. (7 articles), 2004. edited by l. stemmerik & s. stouge. 160.00 6 east greenland caledonides: stratigraphy, structure and geochronology, 93 pp. (6 articles), 2004. edited by a.k. higgins & f. kalsbeek. 160.00 7 review of survey activities 2004, 80 pp. (19 articles), 2005. edited by m. sønderholm & a.k. higgins. 180.00 8 structural analysis of the rubjerg knude glaciotectonic complex, vendsyssel, northern 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weidick & o. bennike. 200.00 15 review of survey activities 2007, 96 pp. (22 articles), 2008. edited by o. bennike & a.k. higgins. 200.00 16 evaluation of the quality, thermal maturity and distribution of potential source rocks in the danish part of the norwegian–danish basin, 66 pp., 2008. by h.i. petersen, l.h. nielsen, j.a. bojesen-koefoed, a. mathiesen, l. kristensen & f. dalhoff. 200.00 17 review of survey activities 2008, 84 pp. (19 articles), 2009. edited by o. bennike, a.a. garde & w.s. watt. 200.00 18 greenland from archaean to quaternary. descriptive text to the 1995 geological map of greenland, 1:2 500 000. 2nd edition, 126 pp., 2009. by n. henriksen, a.k. higgins, f. kalsbeek & t.c.r. pulvertaft. 280.00 19 lithostratigraphy of the cretaceous–paleocene nuussuaq group, nuussuaq basin, west greenland, 171 pp., 2009. by g. dam, g.k. pedersen, m. sønderholm, h.h. midtgaard, l.m. larsen, h. nøhr-hansen & a.k. pedersen. 300.00 20 review of survey activities 2009, 106 pp. (23 articles), 2010. edited by o. bennike, a.a. garde & w.s. watt. 220.00 21 exploration history and place names of northern east greenland, 368 pp., 2010. by a.k. higgins. 200.00 22 lithostratigraphy of the upper oligocene – miocene succession of denmark, 92 pp., 2010. by e.s. rasmussen, k. dybkjær & s. piasecki. 23 review of survey activities 2010, 84 pp. (19 articles), 2011. edited by o. bennike, a.a. garde & w.s. watt. 200.00 24 the east greenland rifted volcanic margin, 96 pp., 2011. by c.k. brooks. 200.00 25 upper cretaceous chalk facies and depositional history recorded in the mona-1 core, mona ridge, 200.00 danish north sea, 60 pp. 2011. by k. anderskouv & f. surlyk. 6363 geological survey of denmark and greenland map series 1 explanatory notes to the geological map of greenland, 1:500 000, humboldt gletscher, sheet 6, 48 pp. + map, 2004. by p.r. dawes. 280.00 2 explanatory notes to the geological map of greenland, 1:500 000, thule, sheet 5 (1991), 97 pp. + map, 2006. by p.r. dawes. 300.00 3 explanatory notes to the geological map of greenland, 1:100 000, ussuit 67 v.2 nord, 40 pp. + map, 2007. by j.a.m. van gool & m. marker. 280.00 4 descriptive text to the geological map of greenland, 1:500 000, dove bugt, sheet 10, 32 pp. + map, 2009. by n. henriksen & a.k. higgins 240.00 5 descriptive text to the geological map of greenland, 1:100 000, kangaatsiaq 68 v.1 syd and ikamiut 68 v.1 nord, 41 pp. + 2 maps, 2010. by a.a. garde & j.a. hollis. 280.00 geology of greenland survey bulletin (discontinued) 181 precambrian geology of the disko bugt region, west greenland, 179 pp. (15 articles), 1999. edited by f. kalsbeek. 240.00 182 vertebrate remains from upper silurian – lower devonian beds of hall land, north greenland, 80 pp., 1999. by h. blom. 120.00 183 review of greenland activities 1998, 81 pp. (10 articles), 1999. edited by a.k. higgins & w.s. watt. 200.00 184 collected research papers: palaeontology, geochronology, geochemistry, 62 pp. (6 articles), 1999. 150.00 185 greenland from archaean to quaternary. descriptive text to the geological map of greenland, 1:2 500 000, 93 pp., 2000. by n. henriksen, a.k. higgins, f. kalsbeek & t.c.r. pulvertaft. replaced by geological survey of denmark and greenland bulletin 18. 186 review of greenland activities 1999, 105 pp. (13 articles), 2000. edited by p.r. dawes & a.k. higgins. 225.00 187 palynology and deposition in the wandel sea basin, eastern north greenland, 101 pp. (6 articles), 2000. edited by l. stemmerik. 160.00 188 the structure of the cretaceous–palaeogene sedimentary-volcanic area of svartenhuk halvø, central west greenland, 40 pp., 2000. by j. gutzon larsen & t.c.r. pulvertaft. 130.00 189 review of greenland activities 2000, 131 pp. (17 articles), 2001. edited by a.k. higgins & k. secher. 160.00 190 the ilímaussaq alkaline complex, south greenland: status of mineralogical research with new results, 167 pp. (19 articles), 2001. edited by h. sørensen. 160.00 191 review of greenland activities 2001, 161 pp. (20 articles), 2002. edited by a.k. higgins, k. secher & m. sønderholm. 200.00 geology of denmark survey bulletin (discontinued) 36 petroleum potential and depositional environments of middle jurassic coals and non-marine deposits, danish central graben, with special reference to the søgne basin, 78 pp., 1998. by h.i. petersen, j. andsbjerg, j.a. bojesen-koefoed, h.p. nytoft & p. rosenberg. 250.00 37 the selandian (paleocene) mollusc fauna from copenhagen, denmark: the poul harder 1920 collection, 85 pp., 2001. by k.i. schnetler. 150.00 prices are in danish kroner exclusive of local taxes, postage and handling note that information on the publications of the former geological survey of denmark and the former geological survey of greenland (amalgamated in 1995 to form the present geological survey of denmark and greenland) can be found on the survey’s website: www.geus.dk title contents abstract introduction setting and stratigraphy materials and methods facies analysis definitions and terminology facies 1: bioturbated chalk facies 2: bioturbated marly chalk facies 3: laminated chalk facies 4: laminated marl facies 5: graded silty chalk facies 6: graded silty chalk with packstone lamination facies 7: graded packstone and wackestone facies 8: non-graded packstone and wackestone facies 9: structureless chalk facies 10: structureless matrix-supported chalk conglomerate facies 11: clast-supported chalk conglomerate facies 12: shear-deformed bioturbated chalk facies 13: shear-deformed matrix-supported chalk conglomerate facies 14: shear-banded chalk slumps genetically related facies facies associations and stratigraphic development facies associations stratigraphic development discussion basin development and depositional history stacking pattern and predictability conclusions acknowledgements references list of publications geological survey of denmark and greenland. bulletin 10, 41-44 ultrabasic alkaline magmatic rocks are products of melts generated deep within or at the base of the lithospheric mantle. the magmas may reach the surface to form lavas and pyroclastic deposits; alternatively they crystallise at depth to form dykes or central complexes. the rocks are chemically distinct and may contain high concentrations of economically interesting minerals and chemical elements, such as diamonds, niobium, tantalum, rare earth elements, phosphorus, iron, uranium, thorium, and zirconium. ultrabasic alkaline rocks are known from several provinces in greenland, but extrusive facies have only been preserved at a few places; e.g. at qassiarsuk in south greenland where pyroclastic rocks occur, and in the maniitsoq region, where a small volcanic breccia (‘fossilik’) contains fragments of palaeozoic limestone. ultramafic lamprophyre and kimberlite are mainly emplaced as dykes, whereas carbonatite forms large intrusive bodies as well as dykes. the ultrabasic alkaline magmas that have been emplaced at certain times during the geological evolution of greenland can be related to major episodes of continental break-up (larsen & rex 1992). the oldest are archaean and the youngest dated so far are palaeogene. figure 1 shows the distribution of known ultrabasic alkaline rocks in west greenland. the large and well-exposed bodies of alkaline rocks and carbonatites in the gardar province were discovered already in the early 1800s (ussing 1912), while less conspicuous bodies were discovered much later during geological mapping and mineral exploration. many alkaline rock bodies, particularly dykes, are difficult to identify in the field because they weather more extensively than the country rock gneisses and form vegetated depressions in the landscape. however, their distinct chemistry and mineralogy render alkaline rocks identifiable in geochemical and geophysical survey data. thus, the sarfartôq carbonatite complex was discovered during regional airborne gamma-spectrometric surveying owing to its elevated uranium and thorium contents (secher 1986). the use of kimberlite indicator minerals has led to the discovery of alkaline rocks such as kimberlites and ultramafic lamprophyres that carry fragments of deep lithospheric mantle. © geus, 2006. geological survey of denmark and greenland bulletin 10, 41–44. available at: www.geus.dk/publications/bull the tikiusaaq carbonatite: a new mesozoic intrusive complex in southern west greenland agnete steenfelt, julie a. hollis and karsten secher 41 fig. 2 100 km 47°w49°w51°w53°w55°w 51°w 49°w 47°w 45°w lamprophyre and kimberlite carbonatite 66°n 68°n 62°n 64°n 60°n 68°n 62°n 64°n 66°n 60°n nuuk kangerluarsoruseq maniitsoq sisimiut tupertalik qaqarssuk tikiusaaq grønnedal-ika qassiarsuk sarfartoq gardar igneous province greenland fig. 1. occurrences of lamprophyre, kimberlite and carbonatite in the southern part of west greenland. such rocks may also contain diamonds. kimberlite indicator minerals are high-pressure varieties of minerals, such as garnet, clinopyroxene, chromite and ilmenite that were formed in the lithospheric mantle. exploration companies have processed thousands of till samples from southern west greenland for kimberlite indicator minerals and found many new dykes. a new carbonatite complex and associated lamprophyre dykes a new carbonatite complex, named tikiusaaq, was discovered in 2005 approximately 100 km east of nuuk, as a result of a field-check of a combined stream sediment geochemical and aeromagnetic anomaly (figs 1, 2). a stream sediment sample collected during the geochemical mapping of west greenland (steenfelt 2001) had a chemical composition that strongly indicated the presence of carbonatitic rocks similar to those at sarfartoq (secher & larsen 1980) and qaqarssuk (knudsen 1991). a small-sized but strong magnetic anomaly (fig. 2) displayed by the aeromagnetic map of west greenland (see e.g. rasmussen 2002) upstream of the anomalous sample also favoured the presence of carbonatite, as large magnetite accumulations are characteristic features of carbonatite complexes. the field work confirmed the presence of the new carbonatite complex. the complex appears to be centred at the position 64°n, 49°46′w, where in situ massive carbonatite was found in the walls of a creek cutting through a gentle slope covered by gravel and vegetation. the preliminary field observations have determined that the emplacement of the carbonatite has affected an area of more than 100 km2 within highly metamorphosed and strongly deformed archaean basement comprising granite, dioritic and tonalitic orthogneiss, amphibolite, and anorthosite. centrally in the area (fig. 2), closely spaced, nearvertical massive carbonatite sheets are up to several metres wide (fig. 3). in their surroundings, the host rocks have been chemically altered (fenitised), and further away the host rocks are fractured and contain numerous small veins (< 1 mm to 20 cm) of carbonatite. the brittle fracturing together with the observation of carbonate-rich breccia indicates that the intrusion was explosive. the dimensions and internal structure of the core zone have not yet been established. however, by analogy with other carbonatite complexes in west green42 5 km lamprophyre dyke or boulder carbonatite core locality outline of carbonatite complex p2o5 % in stream sediment 0.02–0.36 0.37–1.17 1.18–3.95till or stream sediment sample with mantle-derived garnet 50°w 64°n 64°10'n 50°w 64°10'n 64°n nt -276 -242 -220 -203 -189 -171 -156 -143 -132 -121 -110 -96 -83 -68 -49 -27 -2 31 83 fig. 2. close-up of the area, where the new carbonatite was found. the map shows a shaded image of the total magnetic field with large lakes and streams (light blue), and the inland ice (white) superposed on the image. symbols illustrate the concentration of phosphorus (p2o5) in stream sediment and localities where mantle-derived garnet grains have been identified in till or stream sediment. the occurrence of a distinct positive magnetic ano-maly in combination with anomalous stream sediment composition downstream was a strong indication of the presence of alkaline or carbonatitic rocks. the fieldwork confirmed solid carbonatite sheets (dark blue triangle), lamprophyre dykes (black squares), and an area (ringed) of fracturing and veining related to the intrusion of the carbonatite magma. land, it is expected that the core zone is in the order of 1–2 km wide. in addition to the main carbonatite body, many boulders and a few outcropping ultramafic alkaline dykes were discovered in the area north and north-east of the tikiusaaq carbonatite, and these are the likely source of the kimberlite indicator minerals previously recorded in till samples from this area (fig. 2). the observed dykes were only exposed over a few metres, are between 1 and 4 m wide, and have strikes that vary between e–w and ne–sw. mineralogy, chemistry and age carbonatite preliminary field work has identified at least two carbonatite varieties, one pale, creamy white in colour and one brownish grey (fig. 4). the white variety is calcite carbonatite containing disseminated grains or bands of pale green mica, magnetite and apatite. barite and monazite have been found in small amounts. the grey variety is dolomite carbonatite with occasional tiny brown spots suggesting an ankeritic component. boulders of carbonate-rich breccia have abundant xenoliths of the host orthogneiss and garnet-clinopyroxene aggregates (up to 1 cm diameter), along with xenocrysts of phlogopite, olivine, and rare garnet. the matrix of the breccia is similar to that of in situ brown-weathering thin carbonatite dykes within the complex; hence it seems likely that the boulders are derived from a certain phase within the carbonatite. a preliminary age determination of 158 ± 2 ma suggests that the carbonatite is late jurassic. this age was determined by means of u-pb isotope analysis by laser ablation-sector field-inductively coupled plasma mass spectrometry of a zircon crystal from the breccia described above (see also frei et al. 2006, this volume). the age makes the new carbonatite near-coeval with the qaqarssuk carbonatite (173 ma, larsen & rex 1992). further age determinations based on rb-sr isotope chemistry of mica grains from the carbonatite and from one of the lamprophyre dykes are underway. ultramafic dykes boulders and in situ dykes of ultramafic lamprophyres vary in appearance and mineralogy. one widespread type has an olivine-rich groundmass and variable amounts of phlogopite crystals, 2 to 20 mm in size, together with phenocrysts of olivine and magnetite (fig. 5). other occurrences have variable additional carbonate and mica contents in the groundmass. the rocks are possible kimberlites. mantle-derived xenocrysts or xenoliths were not observed macroscopically with certainty, but the presence of mantle-derived minerals will be further tested by means of microscopy and processing of large samples. 43 fig. 3. view northwards along creek with exposed near-vertical sheets of carbonatite emplaced into mafic host rock. intense fracturing with yellow-brown coating of fracture surfaces characterises both the core zone and its surroundings. fig. 4. cream-coloured banded calcitic carbonatite with veins of brown ankeritic carbonatite. the faint banding is caused by variable amounts of disseminated green mica and greyish green apatite. 44 significance of discovery the discovery of the carbonatite in the nuuk region was made in a part of west greenland where alkaline rocks have not hitherto been recorded. the nearest other occurrences are ultramafic lamprophyre dykes recorded at kangerluarsoruseq (formerly færingehavn) on the coast some 100 km westsouth-west of the tikiusaaq carbonatite, and the southernmost of the maniitsoq kimberlites at least 150 km to the north-west. the ultramafic lamprophyre dykes at kangerluarsoruseq yielded ages of 175±7, 185±7 and 196±8 ma (kar dating; larsen & rex 1992), and are thus likely to be part of related intrusive events. carbonatites and kimberlites commonly intrude into weakened zones through the crust, such as large rifts and sutures between former continents. thus, the 560 to 580 ma old sarfartôq carbonatite and associated kimberlites occupy the northern margin of the stable archaean block, and the c. 1300 to 1100 ma old gardar magmas, including the carbonatite within the grønnedal-ika alkaline complex, were emplaced in major rift zones close to the southern margin of the archaean block. the tikiusaaq carbonatite and the nearby lamprophyre dykes, together with the lamprophyre dykes at kangerluarsoruseq are situated near the proposed collision zone between the northern margin of an archaean continent (the tasiusarsuaq terrane) and another continent (friend et al. 1988; mcgregor et al. 1991). this suggests that the newly discovered occurrences of alkaline rocks could be the eastern part of a tract extending along the continent margin, and that more alkaline dykes could be found in this tract. furthermore, the existence of ultramafic lamprophyres in this area opens a possibility of obtaining xenocrystic material from a part of the mantle below greenland that is presently unknown. finally, the discovery of the carbonatite and ultramafic lamprophyres adds to the mineral potential of this part of greenland. references frei, d., hollis, j.a., gerdes, a., karlsson, c., vasquez, p., franz, g., harlov, d., johansson, l. & knudsen, c. 2006: advanced in situ trace element and geochronological microanalysis of geomaterials by laser ablation techniques. geological survey of denmark and greenland bulletin 10, 25–28. friend, c.r.l., nutman, a.p. & mcgregor, v.r. 1988: late archaean terrane accretion in the godthåb region, southern west greenland. nature 335, 535–538. knudsen, c. 1991: petrology, geochemistry and economic geology of the qaqarssuk carbonatite complex, southern west greenland. monograph series of mineral deposits 29, 110 pp. larsen, l.m. & rex, d.c. 1992: a review of the 2500 ma span of alkalineultramafic, potassic and carbonatitic magmatism in west greenland. lithos 28, 367–402. mcgregor, v.r., friend, c.r.l. & nutman, a.p. 1991: the late archaean mobile belt through godthåbsfjord, southern west greenland: a continent–continent collision zone? bulletin of the geological society of denmark 39, 179–197. rasmussen, t.m. 2002: aeromagnetic survey in central west greenland: project aeromag 2001. geology of greenland survey bulletin 191, 67–72. secher, k. & larsen, l.m. 1980: geology and mineralogy of the sarfartôq carbonatite complex, southern west greenland. lithos 13, 199–212. secher, k. 1986: exploration of the sarfartôq carbonatite complex, southern west greenland. rapport grønlands geologiske undersøgelse 128, 89–101. steenfelt, a. 2001: geochemical atlas of greenland – west and south greenland. danmarks og grønlands geologiske undersøgelse rapport 2001/46, 39 pp. (1 cd-rom). ussing, n.v. 1912: the geology of the country around julianehaab, greenland. meddelelser om grønland 38, 1–426. author’s address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ast@geus.dk fig. 5. boulder of ultramafic lamprophyre with olivine rich groundmass and phenocrysts of olivine and magnetite. geological survey of denmark and greenland bulletin 17, 2009, 69-72 69 remote sensing is the science of acquiring, processing, and interpreting images and related data acquired from aircraft and satellites that record the interaction between matter and electromagnetic energy (sabins 1997). the 450–2500 nm wavelength region provides mineralogical information based on analysis of electronic absorption features in transitional metals, especially iron, and of molecular absorption features in carbonate, hydrate and hydroxide minerals (hunt 1977). landsat thematic mapper satellite im ages are widely used to interpret structure and geology, but due to their broad spectral bandpasses landsat images cannot identify specific minerals. however, such details can be achieved by processing and analysing data from hyperspectral sensors. these sensors provide a u nique com bination of high spatial resolution and high spectral resolution imagery of the earth’s surface unavailable from other sources (goetz et al. 1985). an extensive and complex suite of alkaline igneous rocks of carbonatitic and kimberlitic affinity occurs in the baseusing spectral mixture analysis of hyperspectral remote sensing data to map lithology of the sarfartoq carbona tite complex, southern west greenland enton bedini and tapani tukiainen © geus, 2009. geological survey of denmark and greenland bulletin 17, 69–72. available at: www.geus.dk/publications/bull greenland n b ccc d fenite zone outer core of the inner core of the marginal alteration radioactive carbonatite complex carbonatite complex zone shear zones fault51°15´w 66°30´n 2 km 66°n sarfartoq carbonatite complex sukkertoppen iskappe 25 km 51°w kangerlussuaq sø nd re str øm fjo rd inland ice proterozoic archaean sarfartoq carbonatite archaean gneiss reworked in the palaeoproterozoic orthogneiss granitic intrusions thrust a 66°30´n fig. 1. a: geographical and geological position of the sarfartoq carbonatite complex in south ern west greenland (modified from allaart 1982). b: geological map of the sarfartoq carbonatite complex (modified from secher 1986). the rectangle indicates the spatial extent of the hyperspectral data analysed in this study. c: band 6 (510 nm) of the hyper spectral image. d: photograph of the out cropping core zone of the sarfartoq car bonatite within the study area seen from the south-southwest. height of slope is c. 400 m. rosa_2008:rosa-2008 01/07/09 15:48 side 69 70 ment of southern west greenland (lar sen et al. 1983). one of the most important and major carbonatite intrusions is the sarfartoq carbonatite complex (fig. 1). this complex consists of an inner and an outer carbonatite core zone, a fenite zone and a marginal alteration zone (secher & larsen 1980; secher 1986). the objective of an on-going research project of the geological survey of denmark and greenland is to evaluate the use of spectral reflectance techniques and hyperspectral remote sensing data for lithologic mapping and mineral exploration of carbonatites over a known occurrence before applying it to unknown terrains. this short communication reports on the mapping results of spectral mixture analysis of hymap® airborne hyperspectral data covering an important central part of the sarfartoq carbonatite complex (fig. 1). a fuller discussion is presented in bedini (2009). spectral reflectance properties reflectance spectra were acquired using an analytical spectral device (asd inc., usa) field-portable spectrometer, which records 2151 channels within the 350–2500 nm wavelength range. the reflectance spectra of dolomite carbonatites from the sarfartoq carbonatite display characteristic carbonate absorption features with the main co3 2– absorption feature centred around 2320 nm and exhibit a broad ferrous iron absorption feature in the 1000–1300 nm wavelength region (fig. 2). the depth of this broad ferrous absorption feature is positively correlated with the fe2+ content of the dolomite (gaffey 1986). however, in reflectance spectra of ferroan carbonatites with limonitic coating, the broad ferrous absorption feature is diminished due to overlapping spectral reflectance features of ferric iron in limonite. reflectance spectra of carbonatite rocks often display narrow, sharp absorption features at 580 nm, 740 nm, 800 nm and 870 nm, which are attributed to electronic transitions in nd3+ characteristic of rare-earth element-bearing minerals (e.g. rowan et al. 1986). an almost universal characteristic of carbonatite complexes is the presence of a distinctive metasomatic aureole in which the wall rocks have been converted to aegirine-rich and alkali amphibole-rich rocks (secher & larsen 1980). these metasomatic rocks are commonly called fenites and the process fenitisation. reflectance spectra of the fenitic rocks analysed here display an mg–oh doublet absorption feature attributed to the alkali amphibole phase present in fenite. in some cases this is associated with a shallow al–oh absorption feature at around 2200 nm, due to sericite (fig. 2). the marginal alteration zone is distinguished by the hematisation/limonitisation of the country rocks. reflec tance spectra of samples from the marginal alteration zone display intense ferric iron spectral features in the visible and near infrared (vnir) wavelength region (fig. 2; hematised gneiss). this reflectance spectrum in the short-wave infrared wavelength region exhibits an intense al–oh absorption feature at 2200 nm associated with two less intense al–oh absorption features at 2350 and 2450 nm, typical for sericite (fig. 2). sarfartoq hymap data the hymap is an airborne imaging system developed by integrated spectronics, australia, and operated by hyvista corporation. it consists of sensors located on a fixed-wing aircraft typically flown at an altitude of 2.5 km. the sensors collect reflected solar radiation in 126 bands covering the 450–2500 nm wavelength range, including the visible to near-infrared (vnir) and short-wave infrared regions of the electromagnetic spectrum (cocks et al. 1998). the sarfartoq hymap scenes are part of the hypergreen-2002 project of the geological survey of denmark and greenland (tuk iainen & thorning 2005). they were recorded on 9 august 2002 with 4 m nominal pixel size. the hymap data were at mospherically corrected using the atcor4 model (richter & schläpfer 2002). 500 1000 1500 2000 2500 wavelength (nm) r ef le ct an ce in 1 0% in cr em en ts ( of fs et fo r e ac h ro ck t yp e fo r cl ar ity ) hematised gneiss fenite dolomite carbonatite fe2+ o–h fe h–o–h o–h al–oh co h–o–h al–oh mg–oh 3+ co co nd 3 3 3+ 3 2– 2– 2– fig. 2. characteristic reflectance spectra of dolomite carbonatite, fenite and hematised gneiss. rosa_2008:rosa-2008 01/07/09 15:48 side 70 data analysis and results dark pixels (using a threshold of 5% mean reflectance) and green vegetation (using a threshold based on the normalised difference vegetation index, ndvi) were filtered out. the hyperspectral data were mean normalised (i.e. each spectrum was divided by its mean). this form of normalisation eliminates effects of different albedo in the spectral unmixing results (e.g. berman et al. 2004). the minimum noise fraction (mnf) transformation (green et al. 1988) was applied to the normalised data. the mnf is a form of principal components analysis but instead of ordering the data in terms of variance, the data are ordered based on the signal-to-noise ratio. in our case the first 20 mnf bands contained most of the information. image-derived spectral endmembers representing carbonatite, fenite and hematised gneiss were used as input to the spectral mixture analysis (settle & drake 1993), which was applied in the mnf space. the application of the spectral mixture analysis in a subset of mnf bands is advantageous, as noise isolated in the excluded mnf bands does not influence the spectral unmixing (e.g. nielsen 2001). the sum of the fractions was not constrained. the fractions produced from the spectral unmixing analysis were filtered using a 3 × 3 median filter. the spectral unmixing results for the three targets of interest, carbonatite, fenite, and hematised gneiss are shown in fig. 3a–c. a colour composite of the fractions is shown in fig. 3d. the unmixing analysis produced good results for the carbonatite class. the fenite zone and the marginal alteration zone are well mapped within the exposed part of the carbonatite complex along the valley. an important result of the spectral unmixing analysis is the mapping of the outer core zone of the carbonatite consisting of fenitised country rock and carbonatite dykes, distinguished by image analysis as a mixture of fenite and carbonatite (yellow in the colour composite). other parts of the scene are spectrally dominated by the vegetation cover (grass and lichen), although a mixture of gneiss and lichen is also detectable in some parts of the area. field validation of the remote sensing mapping results showed accurate mapping of the carbonatite outcrop. the fenite class shows some confusion with the country rocks. numerous altered spots representing the marginal alteration zone have been detected on the plateau. however, it should 71 fig. 3. fraction abundances produced from the spectral mixture analysis for a: carbonatite. b: fenite. c: hematised gneiss. d: colour composite of the fractions for carbonatite = red, fenite = green and hematised gneiss = blue. yellow is a mixture of fenite and carbonatite. 1 km 1 km 1 km 1 km 51°18’w 51°15’w 51°18’w 51°15’w 51°15’w51°18’w51°15’w51°18’w 66 °3 0’ n 66 °3 0’ n 66°30’n 66°30’n a dc b rosa_2008:rosa-2008 01/07/09 15:48 side 71 be mentioned that the vegetation cover (lichen and grass) reduces its detection. in the arctic environment of southern west greenland green tundra vegetation and lichen constitute a major challenge for remote sensing applied to map surface mineralogy and lithology (e.g. rivard & arvidson 1992). concluding remarks spectral mixture analysis of hymap hyperspectral data was used to map the outcropping rock types (carbonatite, fenite, hematised country rock) of the sarfartoq carbonatite complex. the spectral unmixing produced good results for the carbonatite class, distinguishing the inner and outer core zones (the latter as a mixture of fenite and carbonatite) of the carbonatite complex within the study area. to our knowledge, this is the first study that reports the mapping of fenites using hyperspectral reflectance data. in a hyperspectral re mote sensing study by rowan et al. (1995) of the iron hill carbonatite-alkalic igneous rock complex, colorado, usa, using data from the airborne visible/infrared imaging spectrometer (aviris), fenite could not be distinguished due to low degree of rock outcrop and lower spatial resolution of the hyperspectral data. analysis of high spatial and spectral resolution remote sensing data provides spatially contiguous mineralogical and lithological information for outcropping carbonatite complexes. in inaccessible areas it cannot easily be obtained in any other way. such information is valuable in multi-disciplinary geological studies of carbonatites and, if combined with other types of data obtained by geophysical, geochemical and petrological techniques, it can assist in mapping and mineral exploration of carbonatite complexes. with future high quality hyperspectral data acquired from sensors mounted on satellites (staenz 2009), the availability and areal coverage of such datasets will increase, opening new possibilities for the use of hyperspectral remote sensing in geology. references allaart j.h. 1982: geological map of greenland 1:500 000. sheet 2 frederikshåb isblink – søndre strømfjord. copenhagen: geological survey of greenland. bedini, e. 2009: mapping lithology of the sarfartoq carbonatite complex, southern west greenland, using hymap imaging spectrometer data. remote sensing of environment 113, 1208–1219. berman, m., kiiveri, h., lagerstrom, r., ernst, a., dunne, r. & hun tington j. 2004: ice: a statistical approach to identifying endmembers in hyperspectral images. ieee transactions on geoscience and re mote sensing 42, 2085–2095. cocks, t., jenssen, r., stewart, a., wilson, i. & shields, t. 1998: the hymap airborne hyperspectral sensor: the system, calibration and performance. in: schaepman, m., schläpfer, d. & itten, k.i. (eds): pro ceedings 1st earsel workshop on imaging spectroscopy, zürich, 6–8 october, 1998, 37–43. paris: earsel. gaffey, s. j. 1986: spectral reflectance of carbonate minerals in the visible and near infrared (0.35–2.55 microns): calcite, aragonite, and dolo mite. american mineralogist 71, 151–162. goetz, a.f.h., vane, g., solomon, j.e & rock, b.n. 1985: imaging spectrometry for earth remote sensing. science 228, 1147–1153. green, a.a., berman, m., switzer, p. & craig, m.d. 1988: a transformation for ordering multispectral data in terms of image quality with implications for noise removal. ieee transactions on geoscience and remote sensing 26, 65–74. hunt, g.r. 1977: spectral signatures of particulate minerals in the visible and near infrared. geophysics 42, 501–513. larsen, l.m., rex, d.c. & secher, k. 1983: the age of carbonatites, kimberlites and lamprophyres from southern west greenland: recurrent alkaline magmatism during 2500 million years. lithos 16, 215–221. nielsen, a.a. 2001: spectral mixture analysis: linear and semi-parametric full and iterated partial unmixing in multiand hyperspectral image data. international journal of computer vision 42, 17−37. richter, r. & schläpfer, d. 2002: geo-atmospheric processing of airborne imaging spectrometry data. part 2: atmospheric/topographic correction. international journal of remote sensing 23, 2631–2649. rivard, b. & arvidson, r.e. 1992: utility of imaging spectrometry for lithologic mapping in greenland. photogrammetric engineering & remote sensing 58, 945−949. rowan, l., kingston, m.j. & crowley, j.k. 1986: spectral reflectance of carbonatites and related alkalic igneous rocks: selected samples from four north american localities. economic geology 81, 857–871. rowan, l.c., bowers, t.l., crowley, j.k., anton-pacheco, c., gumiel, p. & kingston, m.j. 1995: analysis of airborne visible-infrared imaging spectrometer (aviris) data of the iron hill, colorado, carbonatitealkalic igneous complex. economic geology 90, 1966–1982. sabins, f.f. 1997: remote sensing: principles and interpretation, 432 pp. new york: w.h. freeman and company. secher, k. 1986: exploration of the sarfartôq carbonatite complex, southern west greenland. in: kalsbeek, f. & watt, w.s. (eds): devel op ments in greenland geology. rapport grønlands geologiske under søgelse 128, 89–101. secher, k. & larsen, l.m. 1980: geology and mineralogy of sarfartôq carbonatite complex, southern west greenland. lithos 13, 199–212. settle, j.j. & drake, n.a. 1993: linear mixing and the estimation of ground cover proportions. international journal of remote sensing 14, 1159–1177. staenz, k. 2009: terrestrial imaging spectroscopy – some future perspectives. 6th earsel sig is workshop, tel aviv, israel, 16–18 march, 2009. tukiainen, t. & thorning, l. 2005: detection of kimberlitic rocks in west greenland using airborne hyperspectral data: the hypergreen 2002 project. geological survey of denmark and greenland bulletin 7, 69–72. 72 authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ebe@geus.dk rosa_2008:rosa-2008 01/07/09 15:48 side 72 geological survey of denmark and greenland bulletin 12, 16-23 16 core samples have been available. as the use of stratigraphic lowest occurrences (lo) of taxa in cuttings samples may be hampered due to downhole caving, the event succession comprises almost exclusively stratigraphic highest occurrences (ho) of taxa (a single significant lo is included in the succession). the event succession is shown in fig. 5a–c; its correlation with international and north sea biozones is shown in fig. 6a–c. seismic sections from the 2-d and 3-d seismic surveys cgd85, dk-1, rtd81–re94, ucg96 and ucge97 have been used to further support the well correlation and to map the stratigraphic units in areas with only scattered well coverage. the combined results from the correlation and mapping procedures are presented as isochore maps for individual stratigraphic units. inspection of cuttings samples from 16 key wells supplemented with sedimentological studies of cored intervals from 23 wells have formed the basis for the lithological and sedimentological descriptions of the units. the well depths mentioned in the lithostratigraphy section are loggers’ depths measured either from rotary table (mdrt) or kelly bushing (mdkb). supplementary data for new type and reference wells are provided in table 1. the names assigned to the new lithostratigraphic units defined herein are derived from nordic mythology and thus follow the nomenclatural tradition previously established for the norwegian north sea (isaksen & tonstad 1989). it should be noted that the micropalaeontology-based palaeoenvironmental terminology used herein was originally developed for a passive margin situation (e.g. the terms ‘neritic’ and ‘bathyal’ to indicate the physiographic zones ‘shelf ’ and ‘shelfslope’, respectively). its application herein to the epicontinental north sea basin solely relates to depositional depth. offshore and onshore lithostratigraphic nomenclature there is a high degree of lithological similarity between the palaeogene–neogene mudstone succession in danish offshore boreholes and that in onshore exposures and boreholes. however, the status of the danish onshore units is quite varied since many units were named before a standard for description of a lithostratigraphic unit was established; some fulfil these requirements, whereas others are still informal. if a previously established onshore unit and an offshore unit can be demonstrated to be identical (e.g. the holmehus formation and the new ve member proposed herein), the name of the onshore unit theoretically has priority over the name of the offshore unit (salvador 1994). in other cases, names of offshore units can be argued to have priority over onshore units (e.g. sele and balder formations over ølst formation). however, in order to acknowledge the traditional distinction between offshore and onshore stratigraphic nomenclature, the two sets of nomenclature are kept separate herein. whenever possible, comments are given in the text to explain the relationship between offshore and onshore danish stratigraphic nomenclature. a correlation between the two sets of nomenclature is shown in fig. 2. chronostratigraphy and biostratigraphy age assessment of the lithostratigraphic units in the north sea sedimentary succession is based on correlation between key biostratigraphic events encountered in the units and the calibrated standard chronostratigraphy published by berggren et al. (1995), with modification for the paleocene–eocene boundary following ratification of its position by the international union of geological scientists (aubry et al. 2002). the key events are from biostratigraphic zonation schemes established for the north sea area. planktonic and benthic microfossils are covered by the zonation schemes of king (1983, 1989; figs 5a–c, 6a–c). dinoflagellates from the paleocene and eocene epochs are covered by the zonation scheme of mudge & bujak (1996b; fig. 6a, b); the oligocene and miocene epochs are covered by the zonation schemes of costa & manum (1988) with modifications by köthe (1990, 2003; fig. 6b, c). key events from these schemes used in this study are listed in fig. 5a–c. for the dinoflagellate events, geochronological calibration has been largely established using age estimates from hardenbol et al. (1998), munsterman & brinkhuis (2004) and williams et al. (2004). for events not mentioned in these three publications, the works of mudge & bujak 17 p2 p9 p7 p6 b a p5 p4 c p8 np13 np12 np10 np9 np11 b a b a p3 c b a p1 pα + p0 np8 np6 np5 np4 np3 np2 np1 np7 abathomphalus mayaroensis cc26 cc25 (pars) pseudotextularia elegans p6 p5 p4 p3 p2 p1 e1a e2a e2b e2c e3a e3b l e1c e1beo ce ne (p ar s) pa le oc en e c re ta ce ou s (p ar s) np14 (pars) e3c 50 60 55 65 nsp6 (pars) nsp5b nsp4 nsb4 (pars) nsb3a nsb2 nsp5a nsb3b nsp3 nsp2 nsp1 nsb1 65.0 ypresian (pars) lo w er (p ar s) u pp er lo w er 55.5 60.0 thanetian selandian danian u pp er (p ar s) maastrichtian (pars) 57.9 54.5 sparnacian c b a b c a planktonic microfossils benthic microfossils dinoflagellate cystsplanktonic microfossils calcareous nannofossils north sea biozonesstandard biozones chronostratigraphy (berggren et al. 1995) stageseries berggren & miller (1988), berggren et al. (1995) mudge & bujak (1996b) martini (1971) king (1989) geochronology ma a fig. 6. biostratigraphic correlation charts showing approximate correlation of calibrated standard planktonic foraminifer and nannofossil biozones with north sea microfossil and dinoflagellate biozones. calibration of the standard biozones follows hardenbol et al. (1998). relationships between the north sea biozones are approximate and their correlation with the standard zones may deviate from that of the original authors (for discussion, see text). a: paleocene–eocene biostratigraphic correlation chart. b: eocene–oligocene biostratigraphic correlation chart. c: oligocene – middle miocene biostratigraphic correlation chart. p17 e6b np14 p9 p7 p8 np13 np12 np11 e2a e2b e2c e3a e3b e1c e1b e3c nsp6 nsp5b nsb4 nsb3ansp5a nsb3b p18 p16 p15 np23 (pars) np22 np21 np19–20 np18 p14 p12 p11 p10 p13 np17 np16 np15 e3d e4a e4b e4c e4d e5a e5b e6a e6c e7a e7b e8b e8a d13 mudge & bujak (1996b) costa & manum (1988), köthe (1990) planktonic microfossils benthic microfossils dinoflagellate cystsplanktonic microfossils calcareous nannofossils north sea biozonesstandard biozones p19 (pars) 35 40 50 45 o lig oc en e (p ar s) lo w er ( pa rs ) eo ce ne ( pa rs ) u pp er m id dl e lo w er ( pa rs ) nsb7ansp9b nsp9a nsb6b nsp8c nsb6a nsp8b nsb5c nsp8a nsp7 nsb5b nsb5a p6 b a np10 (pars) nsp4 (pars) nsb2 (pars) e1a (pars) geochronology ma chronostratigraphy (berggren et al. 1995) stageseries berggren & miller (1988), berggren et al. (1995) costa & manum (1988), köthe (1990), mudge & bujak (1996b) martini (1971) king (1989) rupelian (pars) priabonian 41.3 lutetian ypresian (pars) 49.0 37.0 33.7 bartonian b fig. 6b. eocene–oligocene biostratigraphic correlation chart. 18 m7 m9 m12 m8 m11m10 m6 m5 m3 m2 m4 m1 b a p22 p21 b a p20 p19 p18 nn5 nn6 nn9a– nn7 nn4 nn3 nn2 nn1 np25 np24 np23 np22 np21 (pars) d13 d14 d15 d16 d17 d18 d19 tortonian (pars) burdigalian aquitanian m io ce ne (p ar s) lo w er chattian u pp er rupelian (pars) lo w er (p ar s)o lig oc en e (p ar s) m id dl e langhian serravallian 28.5 23.8 20.5 16.4 14.8 11.2 15 20 30 25 u pp er (p ar s) nsp9a (pars) nsb6b (pars) nsp14b nsb13a nsp14a nsb12c nsp13 nsb12b nsb12a nsp12 nsb11 nsp11 nsp10 nsb10 nsb9 nsp9c nsb8c nsb8b nsb8a nsb7b nsb7a nsp9b m13a (pars) nn9b (pars) planktonic microfossils benthic microfossils dinoflagellate cystsplanktonic microfossils calcareous nannofossils north sea biozonesstandard biozones chronostratigraphy (berggren et al. 1995) stageseries berggren & miller (1988), berggren et al. (1995) costa & manum (1988), köthe (1990), martini (1971) king (1989) geochronology ma c fig. 6c. oligocene – middle miocene biostratigraphic correlation chart. 19 20 (1996b), dybkjær (2004), piasecki (2005) and schiøler (2005) have been consulted. however, whereas hardenbol et al. (1998) and williams et al. (2004) used the timescale of berggren et al. (1995), mudge & bujak used the slightly older timescale from haq et al. (1987) for calibration of their events. therefore, the ages of events only listed by mudge & bujak have been recalibrated herein to conform to the timescale of berggren et al. (1995). king (1989) calibrated his planktonic and benthic microfossil zone markers with the standard chronostratigraphic scale of berggren et al. (1985a, b). however, king noted that only a few first-order correlations were possible; most of the calibrations were made using dinoflagellates, planktonic foraminifers and nannoplankton from onshore sections in the north sea basin (king 1989 p. 420); the correlation of the lower miocene is particularly uncertain (king 1989 p. 446). paleocene and eocene key planktonic and benthic microfossil events from king (1989) were subsequently correlated with the north sea dinoflagellate events by mudge & bujak (1996b). by using the above-mentioned recalibration of key dinoflagellate events from mudge & bujak (1996b), it is feasible to indirectly correlate king’s north sea microfossil events with the timescale of berggren et al. (1995). this has been attempted in fig. 5a–c. figure 6a–c shows the relationships between the north sea biozones and their correlation with the standard planktonic foraminifer and calcareous nannofossil zones. however, it should be noticed that in a few cases the correlation of the north sea microfossil and dinoflagellate zones with the standard zones in fig. 6a–c is at variance with that of the authors of the same zones. this is an effect of improved age determinations of the standard zones and the dinoflagellate events used to calibrate the north sea microfossil zones. the section below outlines the current status for the palaeogene and neogene chronostratigraphic units covered by the studied succession and lists key biostratigraphic events used for chronostratigraphic correlation of the succession. paleocene the bases of the selandian and thanetian stages, which together constitute the upper paleocene series, have yet to be formally defined. however, ongoing work in the international subcommission on palaeogene stratigraphy indicates that the global standard stratotype-section and point (gssp) of the base of the selandian stage will probably be close to the p2–p3a or the p3a–p3b standard planktonic foraminifer zone boundary, while the gssp for the thanetian stage will probably be at the base of magnetochron c26n (gradstein & ogg 2002). hardenbol et al. (1998) followed berggren et al. (1995) in placing the base of the selandian stage at the base of zone p3a, at the lowest occurrence of the planktonic foraminifer morozovella angulata. however, many of the microfossil species that characterise the danian–selandian boundary interval in the international zonation schemes, including m. angulata, are extremely rare or absent in the north sea basin thereby hampering chronostratigraphic correlation of the boundary. based on a study of core material from the type area for the danian and selandian stages, clemmensen & thomsen (2005) concluded that the danian–selandian stage boundary is located in the upper part of the np4 standard nannofossil zone, close to the np4–np5 zone boundary, approximately at the p3a– p3b zone boundary, at c. 60 ma on the timescale of hardenbol et al. (1998). they further concluded that there is a hiatus between the danian and selandian stages in the danish area outside the central graben due to truncation of the danian limestones of the ekofisk formation (fig. 5a; clemmensen & thomsen 2005). hence, the danian–selandian stage boundary is herein placed just below the downhole reappearance (provisional ho) of planktonic foraminifers and the ho of the dinoflagellate alisocysta reticulata, but above the closely spaced events marked by the ho of the planktonic foraminifers subbotina trivialis and globanomalina cf. compressa (e.g. jones 1999; mudge & bujak 2001). the selandian–thanetian stage boundary is herein approximated by the ho of the dinoflagellate palaeoperidinium pyrophorum, at the base of the p5 dinoflagellate zone of mudge & bujak (1996b). this level is close to the base of magnetochron c26n, according to hardenbol et al. (1998). eocene the base of the eocene is at the base of the negative carbon isotope excursion (cie) at 55.5 ma (berggren & aubry 1996; aubry et al. 2002). this position is below the base of the ypresian stage, the lowermost eocene stage. therefore it has been proposed to reintroduce the sparnacian stage as the new basal eocene stage between the cie and the base of the ypresian (aubry et al. 2003). the cie has been correlated with the proliferation of the dinoflagellate genus apectodinium, an event recognised globally (e.g. knox 1996; crouch et al. 2001). onshore denmark, the cie and the proliferation of apectodinium coincides precisely with the laminated stolle klint clay in the 21 2900 m 3000 2900 m 3000 2000 m 2100 2700 m 2900 m horda fm balder fm sele fm lista fm bue mb r og al an d g ro up st ro ns ay g ro up ve mb vile mb våle fm chalk group 3000 2800 kim-1 gr sonic gr sonic gr sonic gr sonic gr sonic kim-1 mona-1 cleo-1 gulnare-1 e-8 e-8 lowermost part of the haslund member of the ølst formation (heilmann-clausen & schmitz 2000; willumsen 2004). in the north sea basin, the acme of apectodinium is located in the lowermost, laminated part of the sele formation (sensu deegan & scull 1977, see below) according to knox (1996). as the event is a lo, its position cannot be determined with certainty in wells in which this interval is covered only by cuttings samples. in the north sea basin, however, this stratigraphic level is characterised by a prominent excursion on the gamma-ray log near the base of the sele formation which therefore can be used as an approximation for the base of the eocene series. the remaining stages of the eocene series, the ypresian, lutetian, bartonian and priabonian stages, lack basal boundary gssps for the present. in this paper, we follow mudge & bujak (1996b) and approximate the bases of the three latter stages by using three key dinoflagellate events: the base of the lutetian stage is at the ho of common eatonicysta ursulae, the base of the bartonian stage is close to the ho of diphyes colligerum, and the base of the priabonian stage is close to the ho of heteraulacacysta porosa. the base of the classic ypresian stage is at the lo of the calcareous nannoplankton species tribrachiatus digitalis. as yet, there is no commonly recognised ho index event at that level in the north sea basin, but the boundary between the sparnacian and the ypresian stages may be placed below the hos of common cerodinium wardenense and apectodinium augustum (fig. 5a), both dinoflagellate species. oligocene the gssp for the eocene–oligocene boundary is in the massignano section (central italy), at the highest occurrence of the planktonic foraminifer genera hantkenina and cribrohantkenina, immediately above the p17–p18 planktonic foraminifer zone boundary (premoli silva & jenkins 1993). however, hantkeninids have not been observed from the north sea basin and alternative zone markers have therefore been used here. in the north sea basin, the planktonic foraminifer globigerinatheka index and the benthic foraminifer cibicidoides truncanus have their hos in the uppermost eocene (king 1989), and the two events may be used to approximate the eocene–oligocene boundary. a palynological marker of the lowermost oligocene is the ho of the dinoflagellate areosphaeridium diktyoplokum (brinkhuis & biffi 1993; brinkhuis & visscher 1995), which is widespread in the north sea basin. the three latter events in combination serve as useful markers for bracketing the eocene–oligocene boundary in the north sea basin. the principal criterion for the rupelian–chattian (lower–upper oligocene) boundary has not yet been defig. 7. log panel illustrating the thickness variation of the rogaland group formations in the danish central graben. 22 cided by the subcommission on palaeogene stratigraphy. indications are that the boundary may be positioned at the base of the p21b planktonic foraminifer zone (premoli silva 2005), at 28.5 ma (hardenbol et al. 1998). however, the defining boundary event cannot be recognised in the north sea basin and its exact correlation with the north sea biostratigraphic event succession remains uncertain. instead, most north sea biostratigraphers recognise the rupelian–chattian stage boundary at the ho of the benthic foraminifer rotaliatina bulimoides. this event marks the top of the nsb7 zone of king (1983, 1989; fig. 5c) and the nsr7 zone of gradstein et al. (1994). the ho of r. bulimoides is at 29 ma in the northern north sea according to gradstein & bäckström (1996), slightly older than the 28.5 ma for the rupelian–chattian stage boundary quoted by hardenbol et al. (1998). the rupelian–chattian stage boundary may also be approximated by the ho of the dinoflagellate rhombodinium draco. in the north sea wells reported herein, where both the hos of r. bulimoides and r. draco have been recorded, these events are largely contemporaneous. however, in the type area of the rupelian and chattian stages, r. draco has its ho above r. bulimoides in the type chattian (van simaeys et al. 2004). therefore, it may be inferred that the two latter events probably bracket the rupelian–chattian boundary (fig. 5c). ve mb tyr mb bue mb log depth core depth våle fm lista fm bor mb gr soni c cecilie-1 clay si. vf. f. m. sand c. vc. p p p p p p p p p ? s 5o 2240 2250 2260 2270 2280 2240 2250 2260 2270 2280 vile mb fig. 8. core log showing intrusive sandstones in the våle and lista formations in the cecilie-1 well. for legend, see fig. 9. the two intervals marked by grey bars in the core depth column are shown as core photographs in fig. 10. 23 lithology sedimentary structures sandstone intrusions mudstone clasts chert siderite calcite concretions trace fossils sandstone mudstone marl chalk carbonate cement (non-calcitic) calcite cement pyrite glaucony parallel lamination faint parallel lamination water-escape pipes (large) load cast dish structures and pipes deformed/slumped bedding fractures/faults bed boundary cross-lamination sandstone intrusions flow structures stylolites zoophycos helminthopsis planolites thalassinoides chondrites low moderate intense degree of bioturbation p g s c miocene the oligocene–miocene series boundary is bracketed by a number of hos at its type section (lemme-carosio, north-west italy). unfortunately, none of the foraminifer events are believed to be true stratigraphic tops (facies dependent), and reworking in the section hampers the use of nannofossil tops (steininger et al. 1997). however, the dinoflagellate succession from the lemme-carosio section has been documented in detail by powell (1986), brinkhuis et al. (1992) and zevenboom (1995, 1996), and provides a means for direct correlation to the north sea basin (munsterman & brinkhuis 2004). the ho of distatodinium biffii is below the chattian–aquitanian boundary in its type section and the ho of chiropteridium spp. is above. this succession of events can be recognised in many north sea wells, and the chattian–aquitanian boundary is positioned between the two. supporting microfossil events that characterise the lowermost miocene include the ho of the diatom aulacodiscus insignis quadrata (small morphotype, same as diatom sp. 3 of king 1983, 1989), a widespread event in the north sea basin, and the ho of the benthic foraminifer brizalina antiqua (king 1989). the ho of the planktonic foraminifer paragloborotalia nana marks uppermost chattian strata. the principal criteria for the aquitanian–burdigalian, burdigalian–langhian and langhian–serravallian stage boundaries are as yet undecided. most authors place the three boundaries at microfossil zone boundaries or magnetochron boundaries at 20.5, 16.4 and 14.8 ma, respectively (hardenbol et al. 1998; williams et al. 2004). the correlation of the three boundaries to the north sea basin is feasible using the dinoflagellate zonation scheme of de verteuil & norris (1996), established for us east coast sections and the review of dinoflagellate index events published by williams et al. (2004). the former zonation scheme is correlated directly with the zonation schemes of berggren et al. (1995) and the miocene timescale by means of calcareous nannofossils and foraminifers. the aquitanian–burdigalian boundary is positioned just above the ho of the dinoflagellate caligodinium amiculum. the burdigalian–langhian boundary is placed between the ho of the dinoflagellates hystrichokolpoma cinctum and pyxidinopsis fairhavenensis, two events that bracket the boundary level. the langhian–serravallian boundary is slightly above the ho of the dinoflagellate cousteaudinium aubryae. in this study, these four events have been used to approximate the three stage boundaries. fig. 9. legend for core logs (figs. 8, 11, 18, 27, 30 and 39); the lithological colour scheme is also adopted on well sections (e.g. fig. 13). review article christiansen & bojesen-koefoed 2021: geus bulletin 47. 6519. https://doi.org/10.34194/geusb.v47.6519 1 of 20 inventory of onshore petroleum seeps and stains in greenland: a web-based gis model flemming g. christiansen*† , jørgen a. bojesen-koefoed geological survey of denmark and greenland (geus), copenhagen, denmark abstract a new inventory on onshore petroleum seeps and stains in greenland has been released by the geological survey of denmark and greenland as a webbased gis model on the greenland mineral resources portal: petroleum seeps and stains in greenland. knowledge on oil and gas seeps, oil stains and solid bitumen occurrences provides key information on mineral and petroleum systems, especially in frontier basins. as the understanding of recent and previous migrations of fluids and gases is important for both mineral and petroleum explorations in greenland, this new inventory has been developed to facilitate exploration and new activities. the classification includes the following types of occurrences: (1) oil seeps, (2) gas seeps, (3) mud diapirs, pingos and gas-rich springs, (4) oil stains in volcanics, carbonates and sandstones, (5) solid macroscopic bitumen and (6) fluid inclusions and other evidence of micro-seepage. the inventory comprises detailed information on localities, coordinates and sample numbers. it also includes descriptions of features and geology, references to data, reports and publications. all information is summarised in either a mineral or petroleum systems context. petroleum seeps and stains have been reported from most palaeozoic, mesozoic and cenozoic basins in greenland where they add important information on petroleum systems, especially distribution and facies variation of source rocks, petroleum generation and later migration, accumulation, remigration, uplift and degradation. the inventory is designed to be updated with additional localities and descriptions and new organic geochemical data. this paper provides a general overview of classification, nomenclature, organisation and content of the inventory. we introduce the regional distribution of petroleum seeps and stains in greenland and general interpretations in the context of mineral and petroleum systems. 1 introduction knowledge on oil and gas seeps, oil stains and solid bitumens provides important information on petroleum systems in most sedimentary basins and may be of great importance in petroleum exploration, especially in frontier areas with no or limited drilling history. in the early part of modern petroleum exploration history, more than a century ago, seepage was an important factor for targeting drilling locations together with surface structural features. most prolific onshore basins in, for example, venezuela, iran, †present address: bygholmvej 15, 2720 vanløse, denmark *correspondence: flemminggc@hotmail. com received: 15 dec 2020 accepted: 09 jul 2021 published: 23 sept 2021 keywords: greenland, organic geochemistry, petroleum and mineral systems, petroleum seeps and stains, solid bitumens abbreviations: bgr: federal institute for geosciences and natural resources casp: cambridge arctic shelf programme gcms: gas chromatography mass spectrometry gcmsms: gas chromatography-tandemmass spectrometry geus: geological survey of denmark and greenland ggu: geological survey of greenland gmom: greenland minerals occurrence map mvt: mississippi valley type deposits sedex: sedimentary exhalative deposits tmax: thermal maturity parameter geus bulletin is an open access, peerreviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: michael b.w. fyhn (geus, denmark) reviewed by: dag arild karlsen (university of oslo, norway) & one anonymous reviewer. funding: see page 16 competing interests: see page 17 additional files: see page 17 https://doi.org/10.34194/geusb.v47.6519 https://orcid.org/0000-0001-6098-9402 https://orcid.org/0000-0001-5647-2769 mailto:flemminggc@hotmail.com mailto:flemminggc@hotmail.com https://creativecommons.org/licenses/by/4.0/deed.ast christiansen & bojesen-koefoed 2021: geus bulletin 47. 6519. https://doi.org/10.34194/geusb.v47.6519 2 of 20 www.geusbul let in.org iraq and central and western us show some surface expressions of petroleum (link 1952; gussow 1954; hunt 1979; macgregor 1993). throughout the last three to four decades, seep materials from all over the world have been analysed and classified by applying modern biomarker analyses (peters et al. 2004 and references therein) and discussed in the context of petroleum systems (abrams 2005; schumacher & abrams 2006). such seep data may add important information concerning distribution and facies variation of source rocks in the subsurface, thermal maturity of source rocks, vertical and lateral migration of generated petroleum and later biodegradation or thermal alteration. the first brief observations on seepage and solid bitumens in greenland go many decades back, see historical comments by, for example, henderson (1969, 1976), mikkelsen (1993) and bojesen-koefoed et al. (2006). systematic field studies combined with modern organic geochemistry began in east and north greenland in the early and mid-1980s and in west greenland in the early 1990s by the present authors and have continued since. at the geological survey of denmark and greenland (geus), the standard routine for most samples collected by geus-geologists has been systematic screening by rock eval pyrolysis. those samples with low tmax (thermal maturity parameter from rock eval pyrolysis) and high production index (a parameter for hydrocarbons already present from rock eval pyrolysis) are then chosen for more detailed organic geochemi stry. christiansen (1994) gave a first review of seeps and other bitumen showings in greenland and provided an overview of the geochemical analyses that were available at that time. seep studies are very important in first evaluations of the petroleum potential of many basins, and in the process of attracting industry interest to greenland exploration (christiansen 2011). understanding the distribution of solid bitumen occurrences as the result of migration of basinal brines or hydrothermal circulation in sedimentary basins is also important for mineral exploration, especially for pb, zn, cu, ag and au. many records of solid bitumen come from the studies of either sedimentary exhalative (sedex) deposits or mississippi valley-type (mvt) deposits (e.g. parnell 1988; jakobsen & omoto 1993; gregg 2004; paradis et al. 2007; gregg & shelton 2012). the role of hydrocarbon compounds during the mineralisation processes is not fully understood in all cases, but likely reflects movements of fluids in faults as known from kongsberg, norway (ag/bitumen), almaden, spain (hg/ bitumen) and cornwall (sn/bitumen). the mineralising deep-basinal brines are often similar in composition and salinity to oil field brines. circulation of hydrothermal fluids with significant amounts of hydrocarbons influences sulphate reduction, and thereby the mineral precipitation in the host rocks, usually fractured dolomite. solid bitumen is also common not only in uraniferous mineralisations (landais 1993, 1996; alexandre & kyser 2006), including various sandstone and conglomerate facies, but also in veins associated with granitic or syenitic intrusions. the role of bitumen and the timing of mineralisation processes are not always clear, but it seems that the interaction between hydrocarbon compounds and uranium-carrying solutions provides good possibilities for co-precipitation of uraniferous mineralised and bitumen under reducing conditions. else, the uranium could precipitate on the surface of previously formed solid bitumen. it is well known that porphyrins in oil carry uranium and other elements like ni and v, as these were incorporated in the paleo-depositional source-rock environment. these metals are generally water soluble in their oxidised state, while they precipitate under reducing conditions. graphite found in mineralisations formed by interaction between organic-rich sediments and magmas is not mentioned in the inventory. some of these occurrences are common in the nuussuaq basin, for example, the graphite andesite in the aaffarsuaq valley on nuussuaq (pedersen et al. 2017). graphite occurrences are common elsewhere in greenland, some associated with organic-rich sources and others clearly abiogenic (see overview in thrane & kalvig 2019). as the understanding of both recent and previous movements of fluids and gases in sedimentary basins in greenland is important in exploration for petroleum and minerals – and for understanding of both minerals and petroleum systems – an inventory of petroleum seeps and stains has been developed at geus to facilitate new exploration models and activities. the inventory, petroleum seeps and stains in greenland, has now  been released in a gis-format on the greenland mineral resources portal and can be accessed at https:// doi.org/10.22008/fk2/jkyxjm. the inventory includes a systematic description of c. 130 localities or small areas across greenland (fig. 1). so far, 52 localities from north greenland are described (selected examples are shown in fig. 2a), 23 in east and north-east greenland (selected examples are shown in fig. 2b), 54 in west greenland (selected examples are shown in fig. 2c; 11 on svartenhuk halvø, one on qeqertarsuaq, three on schades øer, seven on ubekendt ejland, 24 on nuussuaq, two on hareøen and six on disko) and two in south and south-west greenland. for full details of each locality, see the inventory. more than 500 field samples and core pieces have so far been analysed geochemically or described by microscopy. this number is steadily growing. this paper provides a general overview of classification, nomenclature, https://doi.org/10.34194/geusb.v47.6519 http://www.geusbulletin.org https://doi.org/10.22008/fk2/jkyxjm https://doi.org/10.22008/fk2/jkyxjm christiansen & bojesen-koefoed 2021: geus bulletin 47. 6519. https://doi.org/10.34194/geusb.v47.6519 3 of 20 www.geusbul let in.org fig. 1 map of greenland showing distribution of petroleum steeps and stains. red triangles: petroleum seepage and stains associated with source rock. blue circles: petroleum seeps and stains, migrated. white squares: solid bitumen associated with mineral occurrences. yellow stars: gas seepage and gas-rich springs. red text: basins with petroleum seeps or stains including nuussuaq basin, franklinian basin, wandel sea basin, east greenland rift basins, jameson land basin and kangerlussuaq basin. dotted black lines: outline of proterozoic basins or their main outcrop area without petroleum seeps and stains. eb: eriksfjord basin. kb: karrat basin. tb: thule basin. ibf: independence fjord basin. hfb: hagen fjord basin. ebb: eleonora bay basin. f: fossilik. i: ilímaussaq. #*#*#*#* #*#* !( !( #*#* #*#*#* !(!( !(!(!(!(!(!( #*#*#*#*#* #*#*#*#* #*#*#*#*#*#* #*#*#* #* #* #*#* ") ") ") ") ") ")") ") ") !( !(!(!(!(!(!(!(!(!(!(!(!(!(!(!(!(!(!(!(!( !( !(!(!(!(!(!(!(!(!(!(!(!(!(!(!(!(!(!(!(!(!(!(!(!(!( !(!(!(!(!(!(!(!(̂ _̂_ !( #*#* #* #*#* !(#* !( !( !( #* !( #*#* #*#* !( !( ") _̂ _̂ !( !( 400 km davis strait high danmarks havn basin tb franklinian basin wandel sea basin kangerlussuaq basin nuussuaq basin jameson land basin e ast g reenland r ift b asins kb eb i f ebe ibf + hbf 80°n80°n 10°w 10°w 20°w 20°w 30°w40°w50°w 50°w 40°w 60°w 60°w 70°w 70°w 80°w 80°w 0° 0° 70°n 70°n 60°n 60°n https://doi.org/10.34194/geusb.v47.6519 http://www.geusbulletin.org christiansen & bojesen-koefoed 2021: geus bulletin 47. 6519. https://doi.org/10.34194/geusb.v47.6519 4 of 20 www.geusbul let in.org organisation and content of the inventory, and an introduction to the regional distribution of petroleum seeps and stains in greenland with general interpretations in the context of mineral and petroleum systems. 2 classifications and nomenclature seepage features and the various sorts of bitumen occurrences have historically followed different classification systems. some are old and generic (see christiansen 1994), whereas most new studies focus directly on the detailed organic geochemistry, understanding of processes and the implications to the petroleum systems. many countries have petroleum seepage inventories, for example, great britain (selley 1992), italy (burrato et al. 2013), romania (ionescu et al. 2017), china (zheng et al. 2018) and turkey (palabiyik & ozdemir 2019). in greenland, it is necessary to use a modified classification system or nomenclature that can be applied in accordance with the special conditions experienced there. the main reason for this is the arctic climate with d fig. 2 close up examples from the web-based gis model, in this case with a topographic map as background. a: franklinian basin, north greenland encompassing: wulff land – nares land – freuchen land region. b: east greenland rift basins, north-east greenland encompassing: hold with hope region. c: nuussuaq basin, west greenland encompassing: northern disko (base of map) – western nuussuaq – hareøen region. d: text examples (with screen dumps of maps) from the franklinian basin, north greenland (site ng30 on washington land). red triangles: petroleum seepage and stains associated with source rock. blue circles: petroleum seeps and stains, migrated. white squares: solid bitumen associated with mineral occurrences. yellow stars: gas seepage and gas-rich springs. https://doi.org/10.34194/geusb.v47.6519 http://www.geusbulletin.org christiansen & bojesen-koefoed 2021: geus bulletin 47. 6519. https://doi.org/10.34194/geusb.v47.6519 5 of 20 www.geusbul let in.org very cold winter temperatures, short cool summers and, apart from some snow, rather dry conditions. consequently, most onshore sedimentary basins in west and east greenland have several tens of metres thick permafrost – in north greenland, the permafrost is several hundred metres thick. the low temperatures, extensive permafrost, effects of recent glaciations and intense erosion along coasts and in mountain terrains seem to reduce biodegradation of many petroleum seeps or stains. recent glacial erosion has also resulted in modification of outcrops and uplift. organic geochemical analysis is, therefore, a strong exploration tool in greenland compared to regions with warmer and more humid climate, resulting in extensive hydrocarbon degradation. however, knowing that the permafrost may define a secondary seal for shallow oil and gas accumulations and, thus, also impede remigration of oil and gas to the surface, seep data should be interpreted cautiously with respect to defining local drilling targets. in this inventory, we also include information from fully cored boreholes down to a few hundred metres as some of these penetrate sedimentary or volcanic units that often crop out nearby. many of the stratigraphic boreholes from systematic source-rock studies that have been drilled by geus over the years (bojesen-koefoed et al. 2014, 2018, 2020; christiansen et al. 2020a) show petroleum stains. this is also the case for boreholes drilled during mineral or petroleum exploration. whenever possible, we try to distinguish between the following types of occurrences: 1. oil seeps. classical oil seeps are rare in greenland. due to the low temperatures, surface oils have a high viscosity, and in many cases, they are solid or even brittle. many remote seep localities have not been revisited, and it is difficult to evaluate the level of activity. some important localities that have been revisited after initial field work show change from year to year, for example, marraat in the nuussuaq basin. this indicates the presence of active or ‘live’ seeps. however, that activity is probably restricted to a few warm, sunny days (figs 3a–e). 2. gas seeps. this type of seepage is rare if at all present in greenland. most examples of gas are related to gas-rich springs and possibly only with activity in shorter periods of the year (i.e. summer to early autumn). 3. mud diapirs, pingos and other periglacial features and gas-rich springs. pingos and other periglacial features share some characteristics with oiland gas-bearing mud diapirs from prolific petroliferous basins in azerbaijan or romania. pingos are found in many valleys in west and east greenland (bennike 1998). pingos in greenland occasionally show gas bubbles from ‘crater’ lakes or gas-rich water from springs on the side of the pingos (figs 3f, g; see also figs 8a, b in christiansen et al. 2020a). sometimes gas escape structures or small mounds are observed on either muddy or sandy flats (see figs 12a, b in christiansen et al. 2020a). direct evidence of associated petroleum such as iridescence or oil drops on the water surface has not been reported from greenland so far, despite systematic searches. 4. oil stains in volcanics, carbonates and sandstones. this type of petroleum is the most common in greenland. oil stains in volcanics are particularly common in the nuussuaq basin in west greenland (figs 3h–j), whereas oil stains in carbonates are often found in cambrian–silurian strata of the franklinian basin in north greenland, and in the upper permian succession of the east greenland rift basin. oil-stained sandstones are recorded from many different localities and boreholes in north, east and west greenland (see, e.g. fig. 6 in bojesen-koefoed et al. 2020). 5. solid macroscopic bitumen (or pyrobitumen) in mineralised carbonate veins and fractures and in rocks with macro-porosity. this is observed in corals and other fossils, vesicular lavas and partly open fractures (e.g. fig. 6 in christiansen & rolle (1985), plates 7, 8 and 9 in christiansen (1989) and fig. 4a in christiansen & stemmerik (1989)). such types are common in greenland, especially in deeply eroded sedimentary basins with a relatively high thermal maturity such as those found in north and north-east greenland. solid pyrobitumen is also commonly associated with pb-zn-cu mineralisation in veins and fractures in carbonate rocks, especially in north and east greenland (see fig. 5 in jakobsen & omoto (1993)). 6. fluid inclusions and other evidence of microseepage. recognising this type requires a systematic search. hammering on specific healed and mineralised fractures combined with an instant check for petroliferous odour on warm sunny days has disclosed many such localities in the volcanic rocks of the nuussuaq basin in west greenland. most examples have been confirmed later by detailed organic geochemical analyses. such systematic searches have also been applied to outline areas that do not show any apparent evidence of petroleum (i.e. a negative check). in addition to the overall classification, we describe if the petroleum stains and solid bitumen occurrences have an origin from a nearby local source rock that is thermally mature or postmature, with respect to petroleum generation. we describe whether this is due to intrusions of dykes or sills or because the exposed https://doi.org/10.34194/geusb.v47.6519 http://www.geusbulletin.org christiansen & bojesen-koefoed 2021: geus bulletin 47. 6519. https://doi.org/10.34194/geusb.v47.6519 6 of 20 www.geusbul let in.org m l kj i h gf e d cba fig. 3 collection of field photos showing petroleum seepage and stains from greenland. all photos are by the authors unless otherwise specified. dates are indicated as some of the features are known to change over time. a: asphalt seep from portfjeld formation in southern wulff land, north greenland (sample 324200), 18 august 1985. b: asphalt oozing out of partly dolomite filled vug, portfjeld formation in southern wullf land, north greenland (sample 324200), 18 august 1985. c: oil seepage from marraat discovery point, nuussuaq, west greenland, 10 august 2004. d: petroleum oozing out of large vug in basalt from marraat discovery point, nuussuaq, west greenland, 10 august 2004. e: petroleum oozing out of fracture at kuugannguaq valley, disko, west greenland, 14 july 1996. f: water spring with gas seepage from aaffarsuaq valley, nuussuaq, west greenland, 26 july 2019. g: small pools rich in gas on the side of a pingo in the aaffarsuaq valley, nuussuaq, west greenland. blooming algae and vegetation downstream indicate that the water is rich in nutrients, 26 july 2019. h: oil oozing out of freshly broken pillow breccia, nuussuaq, west greenland, 18 july 1996. i: oil oozing out of freshly broken boulder on the beach at sikillingi, nuussuaq, west greenland, 19 july 2006. j: oil oozing out of freshly broken boulder on the beach at sikillingi, nuussuaq, west greenland, 23 july 2006. k: rusty basement rocks with solid bitumen from flade bugt, germania land, 12 august 2009. l: large quartz crystals and solid bitumen (at the point of pencil) in dolomite of the gråklint member in mols bjerge, east greenland, 5 august 1986. m: dolomite with solid bitumen, gråklint member, buch bjerg. https://doi.org/10.34194/geusb.v47.6519 http://www.geusbulletin.org christiansen & bojesen-koefoed 2021: geus bulletin 47. 6519. https://doi.org/10.34194/geusb.v47.6519 7 of 20 www.geusbul let in.org sediments have a regionally high thermal maturity due to subsidence and later uplift. many petroleum stains and solid bitumens are, however, not closely associated with source rocks and are likely to be the result of significant vertical or lateral petroleum migration, typically on a scale from a few kilometres to tens of kilometres. correlation of such petroleum stains and solid bitumens to either known or inferred source rocks is particularly important for understanding the basin history. such correlations can be made by detailed organic geochemistry, either to a known source rock or, in some cases only, to inferred source rocks that are generally interpreted in terms of depositional environment and age. 3 inventory the first step in compiling the inventory was a systematic check of all known publications and reports, and thereby identifying observations, analyses, sample numbers and localities. most field or analytical information has been published or reported internally at geus, especially concerning samples collected during major expeditions and systematic resource evaluations. furthermore, a large in-house sample database at geus, containing >100 000 samples collected by hundreds of geologists over many decades, was checked using search words like pyrobitumen, bitumen, bituminous, petroliferous and oily. in some cases, field geologists with long experience in greenland geology were interviewed. the greenland mineral occurrence map (gmom; accessible at http:// www.greenmin.gl/) provided by geus was also checked systematically for information and descriptions. all samples were assigned to specific localities – or groups of similar localities within shorter distances of some hundred metres, either using aerial photographs, topographic and geological maps of various vintage – and for samples collected since the mid-1990s, directly to gps coordinates. all locations from fieldwork before the mid-1990s – or groups of localities and measured sections – have been georeferenced to modern coordinates by replotting on modern maps or recalculated from various old universal transverse mercator (utm) coordinates or geographical coordinates. in a few cases, georeferencing was done by google maps©. wherever possible, the primary sources of information together with recalculations have been documented. the original measured altitudes were used. most samples from west greenland are close to sea level, so altitude is only indicated for localities higher than 25 m above sea level. some uncertainties remain in the precise identification of coordinates, especially for localities in east and north-east greenland. it is, thus, recommended to check the primary source of information (original maps or photos) in detail, prior to planning any new field work. where possible, the most recent stratigraphical nomenclature is used following watt (2019), but still with reference to the original description when needed for a full understanding. the geographical and basin nomenclature also follows watt (2019). the inventory is documented as a gis compilation, where the user can select to show the distribution of petroleum seeps on the top of standard topographic or geological maps, and other optional data and place names. for each locality, there is a link to the description and classification of features and interpretations. examples of maps exported from the inventory are shown in figs 2a–c, and an example of the text description obtained from the inventory is shown in fig. 2d. where possible, descriptions include photos or references to published photos, references to organic geochemical data, reports and publications. all previous analytical data from geus have been rechecked, not only for quality control and possible contamination but also to update interpretations where more recent data are available. interpretations within a general context of petroleum or mineral systems are provided wherever possible with a special focus on the origin of petroleum and bitumen (source rocks or other organic-rich units). 4 regional distribution of petroleum seeps and stains in greenland petroleum seeps and stains have been reported from most palaeozoic, mesozoic and cenozoic onshore basins in greenland, see regional descriptions in sections 5.1, 6.1 and 7.1, and where possible are presented in stratigraphic order. generally, the seeps and stains provide additional information to the petroleum systems of these basins, especially as a documentation of the generative potential of both known and inferred source rocks. so far, there are no analytically confirmed examples of petroleum seeps or stains from proterozoic basins such as the eriksfjord basin in south greenland, the karrat basin in west greenland, the thule, independence fjord and hagen fjord basins in north greenland and the eleonora bay basin in north-east greenland (fig. 1). this is probably because these basins only have few, and in most cases not known, organic-rich mudstones. furthermore, these basins have a long and complex structural and thermal history with relatively high thermal maturities. thus, if such samples existed, the bitumen would be radically altered. in this context, we do not consider if graphite represents altered bitumen. a few examples of dark bituminous material in fractures have, however, been mentioned from the eleonora bay basin (m. sønderholm, personal communication 2020) in caledonian basement in stauning alper west of https://doi.org/10.34194/geusb.v47.6519 http://www.geusbulletin.org http://www.greenmin.gl/ http://www.greenmin.gl/ christiansen & bojesen-koefoed 2021: geus bulletin 47. 6519. https://doi.org/10.34194/geusb.v47.6519 8 of 20 www.geusbul let in.org jameson land and in intrusions in south greenland (see section 6.1). the origin of this bitumen is uncertain but is most likely related to generation from much younger source rocks. 5 north greenland 5.1 petroleum seeps and stains petroleum seeps and stains are common in the lower palaeozoic franklinian basin in north greenland (chri stiansen 1989; christiansen et al. 1989b). examples have been reported from washington land in the west to peary land in the east (figs 1, 2a), with records from more than 40 localities. the petroleum seeps and stains occur mainly as three types in distinct geological settings (fig. 4): 1. solid bitumen and intense staining or impregnation of carbonates or sandstones closely associated with mature to postmature cambrian source rocks from the henson gletscher formation in southern freuchen land and westernmost peary land (christiansen et al. 1987, 1989b). all examples suggest a maximum vertical migration from the source rock of only a few hundred metres. the organic geoche mical composition and measured reflectance of the bitumen samples are in accordance with the regional thermal maturity of the source rock (christiansen et al. 1987, 1989a; christiansen 1988). 2. soft to solid bitumen closely associated with mature to postmature silurian source rocks in either the lafayette bugt formation or the thors fjord member of the wulff land formation. such bitumen occurrences are widely distributed throughout north greenland from washington land in the west to peary land in the east (christiansen & nøhr-hansen 1989; christiansen et al. 1989b; stemmerik et al. 1997). all examples suggest a maximum vertical migration from the source rock of only a few hundred metres. the organic geochemical composition and measured reflectance of the bitumen are in accordance with the regional thermal maturity of the source rock (christiansen 1988; christiansen et al. 1989a). 3. petroleum seepage and stains in immature cambrian or ordovician strata, especially from the portfjeld and buen formations, in the southern part of warming land and wulff land (figs 3a, b) but possibly also further to the east and west. these examples of migrated oil are all supposed to be the result of long-distance migration from a yet unknown mature lower palaeozoic source rock (christiansen et al. 1989b). in addition to the many such examples described and published by christiansen et al. (1989b), a number of samples have been collected later in western and central peary land (unanalysed or unreported), especially black carbonates and solid bitumens in the henson gletscher formation (c. type 1) and also petroleum stains from older sandstones (c. type 3), and solid bitumens from silurian carbonates of the samuelsen høj formation (c. type 2). furthermore, several carbonate build-ups with pore spaces filled with bitumen have been found in southern pentamerus bjerge and southwards to kap jefferson, washington land, documented as part of the kane project (dawes et al. 2000). samples are still to be analysed, but most likely the bitumen is closely related to silurian source rocks (c. type 2). during the ark-xxv/3 cruise with polarstern in 2010, geologists from the federal institute for geosciences and natural resources (bgr), germany, sampled and subsequently carried out microscopy analysis of bituminous limestones from the lafayette bugt area (kus & pletsch 2012). petroleum seeps and stains are rare in the upper palaeozoic – cenozoic wandel sea basin (fig. 1). this is probably because the organic content of most sedimentary units is low, as only proximal facies crop out. furthermore, regional thermal maturity is either very low (immature with respect to oil generation) or locally very high (postmature related to low metamorphic conditions in some thrust-sheets). a few examples of petroleum stains and solid bitumen have been reported from cores in the triassic dunken formation (dunken-2 borehole) and the upper permian midnatsfjeld formation (kim fjelde-1 borehole) in eastern peary land. both occurrences are found in relatively lean mudstone successions with occasional richer intervals. based on their biological marker signatures, the stains are considered in-situ and intraformational, being generated directly from the rocks in which they are hosted (j.a. bojesen-koefoed unpublished data 2013). 5.2 gas gas seeps have not been recorded from north greenland, and gas was not recorded during core drilling of 13 holes in 1985 (christiansen et al. 1986). from this core drilling, it is known that near surface temperatures are –11 to –14°c, indicating that permafrost in this region is many hundreds of metres thick forming a thick secon dary seal, if shallow gas is present. only a few pingos are known this far north (bennike 1998). gas occurrences have not been systematically examined in relation to oil seeps (i.e. in the form of gas inclusions). 5.3 bitumen associated with mineral occurrences solid bitumen has been found associated with several mineral occurrences in north greenland (fig. 1). one https://doi.org/10.34194/geusb.v47.6519 http://www.geusbulletin.org christiansen & bojesen-koefoed 2021: geus bulletin 47. 6519. https://doi.org/10.34194/geusb.v47.6519 9 of 20 www.geusbul let in.org 390 ma 400 410 420 430 440 450 460 470 480 490 500 510 520 530 540 550 560 570 580 590 600 conglomerate n lithology sandstone mudstone limestone starved basindolostone mound evaporite source rock major petroleum seep minor petroleum stain shallow-water carbonate platform/ramp environmental domains slope deep shelf (carbonate/siliciclastic) siliciclastic shelf (inner, sand-rich) siliciclastic shelf (outer, mud-rich) sand-rich turbidite system muddy turbidite system (including carbonates) navarana fjord lineament harder fjord lineament n h conglomeratic deep-water system n nh h pl s s nn lithology/environment chronostratigraphy central north greenland ( wulff land – peary land) western north greenland (washington land – nyeboe land) lithology/ environment lithostrat lithostrat middle middle o rd ov ic ia n c am br ia n e di ac ar an n eo pr ot er oz oi c lo w er p al ae oz oi c u pp er p al ae oz oi c s ilu ria n d ev on ia n early early furongian epoch 3 epoch 2 terreneuvian pridoli ludlow wenlock late llandovery pl wl mb rg wl mb rg rg db/h ti alal v ? ? ? po pa sk bf bu ti bf bu pf fig. 4 stratigraphy of north greenland (franklinian basin) with simplified relations between petroleum seeps, stains and petroleum systems elements (modified from christiansen 1989 and hopper & ineson in press). al: amundsen land group. bf: brønlund fjord group. bu: buen formation. db/h: dallas bugt and humbolt formations. mb: morris bugt group. pa: paradisfjeld group. pf: portfjeld formation. pl: peary land group. po: polkorridoren group. rg: ryder gletscher group. sk: skagen group. tl: tavsens iskappe group. v: vølvedal group. wl: washington land group. https://doi.org/10.34194/geusb.v47.6519 http://www.geusbulletin.org christiansen & bojesen-koefoed 2021: geus bulletin 47. 6519. https://doi.org/10.34194/geusb.v47.6519 10 of 20 www.geusbul let in.org example is a pb-zn occurrence in the navarana fjord region (east of freuchen land, fig. 2a) that has been described in detail by jakobsen & steenfelt (1985) and jakobsen & omoto (1993). bitumen is associated with several stages of the mineralisation that formed from hot (200°c) metal-bearing brines that migrated through veins and fractures. sulphur isotope analysis suggests that the bitumen correlates with the same source rocks that have generated the long-distance migrated oils in southern warming and wulff land (i.e. a c. type 3; jakobsen & omoto 1993). solid pyrobitumen has also been recorded in other mineral occurrences such as the citronen fjord deposit, peary land (van der stijl & mosher 1998) and in association with zn–pb–ag occurrences on washington land. here, it is found in the ordovician succession close to petermann gletscher in the east and in the silurian succession at kap schuchert in the west (jensen 1998). according to von guttenberg & van der stijl (1993), bitumen has also been noted in zn–pb mineral occurrences in different settings, including silurian reefs at kayser bjerg, hall land, in cambrian carbonate debris flows at hand bugt, northern nyeboe land and at kap wohlgemuth, northern nares land (fig. 2a). based on the examples of mineral occurrences associated with solid bitumen in north greenland, there seems to be a strong correlation between mineral and petroleum systems. some systems are clearly associated with organic-rich silurian mudstones (c. type 2). others are related to organic-rich, thermally postmature, deep-water deposits of cambrian or ordovician age (c. type 3) in strongly folded or steeply dipping zones in the northern part of the franklinian basin. these occurrences are both of mvt and sedex types. 6 east and north-east greenland 6.1 petroleum seeps and stains petroleum seeps and stains are not only common in the rift basins in north-east greenland and in the jameson land basin (fig. 1), especially along the eastern margin on wegener halvø, but also closely associated with exposed jurassic source rocks in central and southern jameson land. given the many organic-rich units in these basins, this is to be expected (fig. 5). it should, however, be noted that several of the many potential source rocks do not have associated bitumen occurrences, probably due to a low thermal maturity in the areas where they crop out. this is the case for the carboniferous lacustrine mudstones (piasecki et al. 1990; stemmerik et al. 1990), jurassic resinite-rich coals (bojesen-koefoed et al. 1996; petersen et al. 2002, 2013) and the jurassic sortehat member (krabbe et al. 1994), where associated bitumen is still not found. in total, 19 localities or small areas with petroleum seeps and stains have been reported so far. petroleum stains were discovered in a fault-zone in the metamorphic basement close to flade bugt on germania land by henrik stendal and svend monrad jensen in 1990 (fig. 3k). preliminary organic geochemical studies showed that the petroleum stains are severely degraded, with a most likely origin from a saline lacustrine unit (christiansen et al. 1991). more recent sampling followed by gas chromatography-tandem-mass spectrometry (gcmsms) indicates that the oil was probably generated from marine shale source rocks of jurassic – lower cretaceous age (bernbjerg formation), broadly equivalent to the principal source rocks of the north sea graben systems (bojesen-koefoed et al. 2009). the stains may represent long-distance migration from the offshore danmarkshavn basin or generation from an ancient sedimentary cover, which has later been removed by erosion. organic-rich lacustrine mudstones of potential source-rock quality occur in two intervals in the devonian succession: the middle devonian (givetian) and the upper devonian (?famennian; christiansen et al. 1990a). in both cases, black solid bitumen in calcite veins is associated with the source rocks. solid bitumen associated with devonian lacustrine units on canning land has been mentioned by the cambridge arctic shelf programme (casp). oil stains have been reported from a succession overlying the upper permian ravnefjeld formation at kap stosch in the north-western corner of home forland (fig. 2b), within both cretaceous sandstones and paleocene basalts and sediments. one sample of basalt carries tar-like material in vesicles and another sample shows petroleum hosted by calcite veins in basalt. biological marker data clearly link both occurrences to the upper permian ravnefjeld formation, which crops out in the area (bojesen-koefoed et al. 2009). the upper permian carbonates on gauss halvø contain black solid bitumen, and overlying triassic sandstones are stained by oil; in both cases, unpublished biological marker data suggest the upper permian ravnefjeld formation as the source. solid black bitumen is common in fractures and vugs in the upper permian carbonates and lower triassic sandstones on wegener halvø (stemmerik et al. 1989; scholle et al. 1991). these bitumen occurrences have been generated from the ravnefjeld formation source rock that has a high thermal maturity in this region (ro: 1.4–1.8%; christiansen et al. 1990b). in contrast, the thermal maturity of the exposed ravnefjeld formation is low along the western margin of the jameson land basin in schuchert dal, and no records of solid bitumen or oil stains have been made, neither in the field nor from the numerous core holes in this area. highly reflecting solid bitumen (ro ≥ 2.5%) has been reported from vugs in a dolomite-calcite-quartz vein https://doi.org/10.34194/geusb.v47.6519 http://www.geusbulletin.org christiansen & bojesen-koefoed 2021: geus bulletin 47. 6519. https://doi.org/10.34194/geusb.v47.6519 11 of 20 www.geusbul let in.org jameson land base-of-slope fans channel fills marine mudstone marine carbonate deep marine sandstone (beyond shelf break) shallow marine sandstone (shelf) humid terrestrial (including lacustrine) arid terrestrial sedimentary environments extrusive igneous rocks lithological symbols other symbols anhydrite/gypsum coal layers sandstone siltstone mudstone/shale organic rich mudstone/shale limestone prograding shelf-margin sediment wedge lavas: mafic conglomerate dolomite bioherm wollaston foreland – kuhn ø hochstetter forland – stormlandet ? ? traill ø hold with hope ? ? ? ? early middle late early middle late paleocene cisuralian guadalupian lopingian early middle late early middle late early late eocene oligocene miocene pliocene pleistocene mississippian late mid early pennsylvanian 259.8 tournaisian famennian frasnian givetian eifelian emsian pragian lochkovian visean serpukhovian bashkirian moscovian kasimovian gzhelian asselian sakmarian artinskian kungurian roadian wordian capitanian wuchiapingian changhsingian anisian ladinian carnian norian rhaetian hettangian sinemurian pliensbachian toarcian aalenian bajocian bathonian callovian oxfordian kimmeridgian tithonian berriasian valanginian hauterivian barremian aptian albian cenomanian turonian coniacian santonian campanian maastrichtian danian selandian thanetian ypresian lutetian bartonian priabonian blosseville home forland brorson halvø hall bredning vardekløft kap stewart scoresby land foldvik seepage in basement traill ø celsius bjerg kap kolthoff vilddal scoresby sund wollaston forland rupelian chattian aquitanian burdigalian langhian serravallian tortonian messinian piacenzian/zanclean lt. pleist./gelasian era pa le oz oi c m es oz oi c ce no zo ic age period epoch age/stage 0 10 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 210 220 230 240 250 260 270 280 290 300 310 320 330 340 350 360 370 380 390 400 410 ca rb on ife ro us de vo ni an pe rm ia n tr ia ss ic ju ra ss ic cr et ac eo us pa le og en e ne og en e 358.9 372.2 382.7 387.7 393.3 407.6 410.8 419.2 346.7 330.9 323.2 315.2 307.0 303.7 298.9 295.5 290.1 279.3 272.3 268.8 265.1 254.2 252.2 250.0 247.1 241.5 237.0 228.4 209.5 201.3 199.3 190.8 182.7 174.2 170.3 168.3 166.1 163.5 157.3 152.1 145.0 139.4 133.9 130.8 126.3 113.0 100.5 93.9 89.8 86.3 83.6 72.1 66.0 61.6 59.2 56.0 47.8 41.2 37.8 33.9 28.1 23.03 20.44 15.97 13.82 11.63 7.25 ryazanian volgian induan olenekian te rtia ry quat. chronostratigraphy group ? source rock borehole with oil major oil seep minor oil fig. 5 stratigraphy of east and north-east greenland with simplified relations between petroleum seeps, stains and petroleum systems elements (modified from christiansen et al. 1992; fyhn et al. 2021). in a dolomite bed from the triassic gråklint member, mols bjerge on traill ø (marcussen et al. 1987; fig. 3l). asphalt has also been found within fractures of the gråklint member and the solfaldsdal formation on buch bjerg close to carlsberg fjord (thomassen 1973; fig. 3m). sandstones associated with the lacustrine source rocks in the upper triassic – lower jurassic kap stewart https://doi.org/10.34194/geusb.v47.6519 http://www.geusbulletin.org christiansen & bojesen-koefoed 2021: geus bulletin 47. 6519. https://doi.org/10.34194/geusb.v47.6519 12 of 20 www.geusbul let in.org group in northern jameson land (dam & christiansen 1990) occasionally contain highly coalified bitumen near palaeogene sills. dark triassic and jurassic sandstones with pyrobitumen have been reported at several places on traill ø and geographical society ø. they were described first as evidence of a paleo oil field at laplace bjerg (marcussen et al. 1987) and later as large exhumed oil fields at svinhufvud bjerge, mols bjerge, bjørnedal and laplace bjerg (price & whitham 1997; andrews et al. 2020a, 2020b). due to limited analytical data, very high thermal maturity (christiansen & boserup 1990) and lack of non-degraded – and analytically confirmed – petroleum stains in the vicinity of these features, the model is considered controversial and has been challenged by many oil companies and geus geologists (christiansen 2011; christiansen et al. 2020b). however, as some solid bitumen has been recorded – and to make the inventory as comprehensive as possible and, importantly, to advocate for further studies – these localities have been included in the detailed locality descriptions. oil-stained sandstones occur in several of shallow boreholes drilled by the geological survey of greenland (ggu) in the sjællandselv area of southern jameson land, where the upper jurassic hareelv formation source rock has been penetrated by shallow-core drilling (piasecki 1987; christiansen & boserup 1990). the cores were studied in more detail for their organic geochemistry by requejo et al. (1989). the same sourcerock interval was drilled in 2008 (the blokelv-1 core), and non-degraded crude oil was noted bleeding from a belemnite (see fig. 8 in bojesen-koefoed et al. 2018) as well as from carbonate-lined fractures in a dolerite sill. all these oils have a composition that can be directly correlated to the upper jurassic hareelv formation source rock. exposed oil-stained sandstones have so far not been reported in this part of jameson land but are likely to occur. jurassic sandstones of a similar age on milne land, west of jameson land, have a locally petroliferous odeur (m. larsen, personal communication 2020). water from streams in the same area was described as unfit for drinking and for brewing tea or coffee due to natural petroleum contamination (c. heinberg, personal communication 1984). perregaard & schiener (1979) report some effects of the intrusion of a dyke into kimmeridgian organic-rich claystones on milne land, and although oil staining is not mentioned, its presence seems implicit in their findings. their samples were originally collected in 1977 by s. piasecki, who, however, does not recall if staining was present and does not have any notes concerning these matters in his diary. recording potential traces of petroleum was not a priority at the time, meaning that the presence of such material cannot be ruled out (s. piasecki, personal communication 2020). the nanok-1 fully cored borehole at hold with hope (fig. 2b) penetrated a thick cenomanian–campanian succession (bojesen-koefoed et al. 2020). impressive oilstained sandstones (see fig. 6 in bojesen-koefoed et al. 2020) occur at several intervals, and a detailed organic geochemistry suggests the presence of an underlying marine mid-cretaceous source rock. similar oil-stained sandstones are exposed at distances >1 km from the borehole, indicating that this is a semi-regional feature. pale, oozing material has been reported from partially agate-filled cavities in palaeogene basalt on hold with hope (w. s. watt, personal communication 2008, 2020). unfortunately, both the sample and extract were lost after a laboratory relocation, following a preliminary description by f. g. christiansen and j. a. bojesen-koefoed. tar-like material has been reported from a palaeogene basalt south of scoresby sund close to a major fault zone (watt & wrang 1984). the material is strongly degraded, but δ13c-isotope and pyrolysis-gc data are in accordance with generation from an upper jurassic source rock (requejo et al. 1989). petroleum seeps or stains have not been reported from the cretaceous–palaeogene kangerlussuaq basin, southern east greenland. however, some indications of hydrocarbons from fluid inclusions in sandstones suggest local generation and migration (jonk et al. 2005). 6.2 gas gas seeps have not been recorded in north-east and northern east greenland. due to the very thick permafrost in this region, pingos mainly occur in the southern part, especially on geographical society ø and traill ø (bennike 1998). the tobias dal, home forland, hosts a few pingos. pingos are common in the jameson land basin (bennike 1998). one poorly documented gas seep in colorado dal, northern jameson land, found and reported by the company arco during their exploration in the 1980s is probably from a pingo (see historic photo in bennike 1998). gas kicks or free gas have not been recorded from any of the 51 fully cored boreholes drilled by ggu and geus in the 1980s, or more recently between 2008 and 2017. gas from canned core pieces has been systematically analysed from many drilling locations in both east and north-east greenland (e.g. karlsen et al. 1988). gas has been observed, sampled and analysed from a hot spring in rømer fjord south of scoresby sund in the 1970s (w.s. watt, personal communication 2020), which may suggest that the mesozoic succession is present in the subsurface much farther south than jameson land. https://doi.org/10.34194/geusb.v47.6519 http://www.geusbulletin.org christiansen & bojesen-koefoed 2021: geus bulletin 47. 6519. https://doi.org/10.34194/geusb.v47.6519 13 of 20 www.geusbul let in.org 6.3 bitumen associated with mineral occurrences mineralisation with bitumen in east greenland has been recorded from some of the same geological settings that are important for the main elements of petroleum systems, for example, stratiform cu-pb-zn mineralisation in the organic-rich mudstones of the ravnefjeld formation, previously known as the posedonia shale, equivalent to the kupferschiefer in europe (thomassen 1973; harpøth et al. 1986). stratiform pb-zn-cu mineralisation also occurs within the thin organic-rich dolomite units of the triassic gråklint member, associated with solid bitumen (thomassen et al. 1982). bitumen is also associated with cu-zn-pb stratabound mineralisations in permian carbonate build-ups, especially on wegener halvø (thomassen et al. 1982; harpøth et al. 1986). uraniferous hydrocarbons, also called carburan using the french terminology, have been reported from devonian acid volcanic rocks from randbøldal, gauss halvø (secher & steenfeldt 1976; secher et al. 1976). the carburan is found as globular grains and veinlets in cavities in the rhyolites close to faults and fracture systems, and with a close association between bitumen and uranium (a. steenfeldt, personal communication 2020; k. secher, personal communication 2020). the source is not documented for neither the hydrocarbons nor the uranium, albeit porphyrins in oil can carry uranium and many source rocks are enriched in uranium. 7 central west greenland 7.1 petroleum seeps and stains petroleum seeps and stains are common in the volcanic rocks of the nuussuaq basin in west greenland ( christiansen et al. 1996b; bojesen-koefoed et al. 1999; christiansen et al. 2020a; figs 1, 2c). however, before the first non-degraded oil stains were found in 1992 and further substantiated by drilling the year after at marraat on western nuussuaq (christiansen 1993; christiansen et al. 1994a, 1994b, 1996b), only few scattered observations of solid bitumen had been reported by, for examp le, henderson (1969, 1976) and pedersen (1986). since the breakthrough with drilling of the marraat-1 well in 1993, systematic search for petroleum stains and subsequent organic geochemical analysis has been a standard procedure during many geus expeditions, and hundreds of samples have been collected and analysed in detail. petroleum stains occur in a large area, at least 50 km by 150 km, including parts of disko, hareøen, western nuussuaq, ubekendt ejland, schades øer and svartenhuk halvø (bojesen-koefoed et al. 1999; pedersen et al. 2008; christiansen 2011; christiansen et al. 2020a; figs 1, 2c). petroleum exploration by a small canadian company grønarctic energy was partly driven by this new knowledge on petroleum seeps. many of their drilled cores show evidence of oil or gas (christiansen et al. 1995, 1996a; bojesen-koefoed et al. 1997, 1999). as a result of this large and systematic oil hunting effort, five distinct oil types have been documented in detail and correlated to known or inferred source rocks (christiansen et al. 1996a, 1996b, 2020a; bojesen-koefoed et al. 1997, 1999, 2004, 2007; pedersen et al. 2006; sørensen et al. 2017; fig. 6): 1. marraat oil type. high-wax oil with characteristic angiosperm-derived biomarkers generated from a paleocene source rock in the eqalulik formation. intense seepage and stains in a well-defined, more than 2000 km2 area, on western nuussuaq, hareøen and disko, between the itilli and kuugannguaq– qunnilik fault zones. more than 60 samples from 25 localities including the marraat-1, ganw#1, gane#1 and gank#1 core holes have been analysed and documented. 2. kuugannguaq oil type. high-wax oil generated from an inferred terrigenous cretaceous source rock. this oil type was originally restricted to a small area on northern disko close to or within the kuugannguaq fault zone. furthermore, a number of samples collected in the ubekendt ejland – svartenhuk halvø area have been referred as to the kuugannguaq oil type, mainly mixed with the itilli oil type. more than 30 samples have been analysed and documented. 3. itilli oil type. low-wax oil generated from inferred mid-cretaceous marine source rocks with some facies variations. this oil type is distributed throughout the nuussuaq basin and is the dominating type not only in the northern and westernmost parts including ubekendt ejland, svartenhuk halvø and nuussuaq west of the itilli fault zone but also at asuk on disko. in most cases, this oil type is only found as microseepage in mineralised veins in volcanics of the vaigat formation. a biodegraded subtype of the itilli oil occurs in oil-stained sandstones at asuk on disko. more than 60 samples from at least 30 localities have been analysed and documented. these numbers include some examples of mixing with both the marraat, kuugannguaq and eqalulik oil types. 4. eqalulik oil type. this oil type is very characteristic in composition. it was originally assumed to have been generated from a local lacustrine source rock of early cretaceous age, but unpublished data have made this assumption dubious. with one exception, this oil type is only found as mixtures with the marraat oil type on nuussuaq or with the itilli oil type on svartenhuk halvø. distribution on nuussuaq is restricted to a minor area along the vaigat coast, and https://doi.org/10.34194/geusb.v47.6519 http://www.geusbulletin.org christiansen & bojesen-koefoed 2021: geus bulletin 47. 6519. https://doi.org/10.34194/geusb.v47.6519 14 of 20 www.geusbul let in.org fig. 6 stratigraphy of west greenland (nuussuaq basin) with simplified relations between petroleum seeps, stains and petroleum systems elements (based on bojesen-koefoed et al. 1999; dam et al. 2009; pedersen et al. 2017; sørensen et al. 2017; and slightly modified from christiansen et al. 2020a). an: anaanaa member. i: itilli type. k: kuugannguaq type. m: marraat type. n: niaqornaarsuk type. na: naujánguit member. or: ordlingassoq member. upper cretaceous paleocene eocene oligocene miocene pliocene quaternary lower cretaceous 0 ma 5 10 15 20 25 30 35 40 45 50 55 60 65 70 75 80 85 90 95 100 105 110 115 120 aptian albian cenomanian turonian coniacian santonian campanian maastrichtian danian selandian thanetian ypresian lutetian bartonian priabonian rupelian chattian aquitanian burdigalian langhian serravallian tortonian mainly continental deposits mainly marine or deltaic sandy/silty deposits, locally mudstones mainly marine mudstones, locally sandy/silty messinian zanclean piacenzian p al eo ce ne e oc en e 61 –6 0 60 –5 8 56 –5 4 54 –5 3 39 –3 8 62 –6 1 m al ig aâ t f m s va rte nh uk f m n aq er lo q fm e rq ua fm h ar eø en fm va ig at f m ma volcanic stratigraphy chronostratigraphy sedimentary stratigraphy nw se nw se source rock unconformity (hareøen) (hareøen) kangilia fm itilli fm itilli fm atane fm agatdal fm eqalulik fm m n k i quikavsak fm lake, pingo with gas core hole with gas core hole with oil core hole with oil and gas major oil seep minor oil stain an or na in the gane#1 and gank#1 cores. more than 40 samples with an eqalulik oil-type fingerprint have been analysed and documented. 5. niaqornaarsuk oil type. high-wax oil generated from campanian mudstones. these oil stains have only been reported from two small areas close to the qunnilik fault zone along the vaigat coast and in the aaffarsuaq valley, apparently without mixing with other oil types. more than 10 samples from six localities have been analysed and documented. https://doi.org/10.34194/geusb.v47.6519 http://www.geusbulletin.org christiansen & bojesen-koefoed 2021: geus bulletin 47. 6519. https://doi.org/10.34194/geusb.v47.6519 15 of 20 www.geusbul let in.org this well-defined pattern of petroleum stains with distinct oil types and some systematic evidence of mixing provide strong indications on the distribution, facies variation and thermal maturity of the source rocks in the subsurface, and of major structural elements. despite the large number of analysed samples from many different localities, there are still unresolved questions, especially related to details on facies variations of source rocks, how and where the different oil types migrated and mixed, and the lateral and vertical migration di stances (christiansen et al. 2020a). 7.2 gas pingos are common in the nuussuaq basin (bennike 1998). gas seepage, gas-rich springs or lakes with bubbling gas have been in described several places over many decades (see christiansen et al. 2020a). gas has also been recorded many times during core drilling by geus or mineral exploration companies. these observations are only included in the inventory if there is direct evidence that gas reached the surface (see christiansen et al. 2020a for an overview, locations and geochemistry). for this reason, some gas observations from serfat, vismann core holes and umiivik-1 (dam et al. 1998) are not included in the inventory, whereas gas associated with petroleum stains in marraat-1, ganw1, gane#1, gank#1 and gant#1 is described together with details of the organic geochemistry of the different oil types. although the analysed gases are comparatively dry, most data indicate a clear thermogenic fingerprint. some of the sampled gases from drill holes are directly associated with specific oil stains in the penetrated sediments or volcanics, especially related to the marraat and eqalulik oil types (christiansen et al. 2020a). 7.3 bitumen associated with mineral occurrences solid pyrobitumen has been recorded in epithermal sulphide and gold mineralisations on the southern part of ubekendt ejland (bernstein & knudsen 2004). these mineralisations in veins of the volcanic succession are relatively close to areas with non-degraded petroleum stains of the itilli oil type or mixture with the kuugannguaq oil types. 8 south and southern west greenland minor traces of petroleum and hydrocarbon-rich fluid inclusions have been reported from the ilímaussaq intrusion in south greenland (fig. 1; e.g. petersilie & sørensen 1970; konnerup-madsen et al. 1979). the hydrocarbon compounds were interpreted as abiogenic until recently, when detailed organic geochemistry of extracts from the most common rock types naujaite, kakortokite and lujavrite has shown specific biomar kers characteristic of upper cretaceous and paleocene source rocks farther north in west greenland (laier & nytoft 2012a, 2012b). this suggests that the petroleum recorded originated via long-distance migration during the palaeogene from offshore sedimentary basins or hydrothermal circulation from previously overlying sediments. the fossilik sedimentary inlier, situated in the basement area north of nuuk (fig. 1), was originally interpreted as a downfaulted erosional relict (poulsen 1966), but later reassessment showed it to represent a fallback breccia sitting in a jurassic-age diatreme (larsen et al. 2009). the breccia includes, among many other lithologies, a grey carbonate of late ordovician age with small droplets of oil. the oil is non-degraded and, based on biological marker data, is known to have been generated from a carbonate source rock (bojesen-koefoed 2011). the precise origin of this oil is not known, since no deposits having petroleum source potential are found in the fossilik inlier, but it is most likely intraformational. the composition of the fossilik oil is unique and shows little compositional similarity to any known ordovician source rock or oil in the arctic region. 9 offshore the offshore basins of greenland have not been included in this inventory. still, both conventional and shallow seismic data indicate that offshore seepage could be common in many basins. numerous examples of possible flat and bright spots, gas chimneys, gas hydrates and pockmarks have been described in many papers, including also so-called satellite slicks indicative of seepage (e.g. fig. 4c in bojesen-koefoed et al. 2007; numerous figures in cox et al. 2021; and figs. 3, 4 in gregersen et al. 2007). however, so far neither seeping oil nor gas has been sampled and analysed from the seabed, the water column or at the water surface. an oil-stained ordovician-age lump of carbonate was recovered by dredging at the davis strait high (dalhoff et al. 2006). the oil shows a relatively wax-rich distribution of normal alkanes, with pristane/phytane ratio less than unity. however, biological marker data do not show obvious carbonate characteristics and are dissimilar to those of the fossilik oil stain (see section 8). it is more than likely that future research into the offshore domain will result in numerous manifestations of petroleum migration. 10 conclusions and recommendations petroleum seeps and stains are widespread in most palaeozoic–mesozoic basins in greenland. in fact, the https://doi.org/10.34194/geusb.v47.6519 http://www.geusbulletin.org christiansen & bojesen-koefoed 2021: geus bulletin 47. 6519. https://doi.org/10.34194/geusb.v47.6519 16 of 20 www.geusbul let in.org number of occurrences is striking; bearing in mind that, historically, geological mapping did not necessa rily focus on manifestations of seeps, and given the permafrost conditions, which do not facilitate petroleum movement. the seeps and stains documented here, thus, reflect the ubiquity of petroleum in the sedimentary basins of greenland. the overview presented, thus, provides important information of past and present petroleum systems, especially concerning distribution, facies variation and thermal maturity of known or inferred source rocks. this review also provides information about migration, trapping and later degradation history. this provides important input for resource evaluations and input to new exploration models. documentation of basinal brines and hydrothermal circulation with petroleum compounds is also important for some mineralisation models. this is especially relevant for mvt or sedex occurrences in palaeozoic basins, and stratiform mineralisation associated with organic-rich units. considering the size of greenland, and the rather limited number of historic field campaigns, there is no doubt that many more petroleum seeps and stains remain to be discovered. this emphasises the old explorer’s rule of thumb that ‘you find what you look for’. however, the chance for finding new petroleum seeps or stains is much better today than previously. compared to the systematic fieldwork in the 1980s and the 1990s, it is now possible to use modern satellite and airborne hyperspectral sensing tools to generate data with good proxies for petroleum seeps. in addition, advanced 3d-geological techniques can focus systematic ground control on localities that fulfil specific geological criteria or display anomalous features, which should be exami ned in more detail. we see the new inventory as a modern dynamic tool that can be applied in future mineral and petroleum explorations. it also provides background documentation for research on fluid and gas movements in sedimentary basins. furthermore, it is likely that more gas seepage will be observed in the coming years due to permafrost degradation, which, in some cases, create specific and rare ecosystems. the first version of this inventory was based on field information from a limited number of geus geologists and subsequent analytical work on their samples at geus laboratories. we are confident that other research groups or exploration teams may have additional observations on petroleum seeps and stains in greenland, with or without analytical evidence. we invite geologists to forward information on localities, samples, features and geological settings together with key data, references and any preserved sample material for inclusion in future versions of the inventory. it is important that such material could be systematically analysed using similar methods as previous studies by geus geologists, including new state-of-the-art methods where possible. knowing that the analytical work has been carried out over several decades and acknowledging the ongoing development and sophistication of analytical techniques and interpretation schemes, some critical samples should be reanalysed with modern techniques. for example, gcmsms and compound specific isotope analyses could provide better oil-to-oil and oilto-source rock correlations. this will provide a better understanding of the distribution patterns of past and present mineral and petroleum systems. the samples collected previously and the extracts analysed at geus are, with the few exceptions where only limited material was available for a one-shot analysis, still available for future studies and applications of new methods that can provide further details on petroleum and mineral systems. acknowledgements first of all thanks to all those dedicated field geologists who, on purpose or thanks to their curiosity, have made observations on petroleum seeps and stains in greenland: finn dalhoff, jon ineson, ulla h. jacobsen, hans f. jepsen, olav nykjær, john s. peel, thomas pletsch, anders pilgaard, flemming rolle, diogo rosa, agnethe steenfelt, lars stemmerik and martin sønderholm in north greenland; morten bjerager, gregers dam, svend monrad jensen, mikael larsen, poul-henrik larsen, henrik olsen, stefan piasecki, anders pilgaard, karsten secher, lars stemmerik, henrik stendal, jens therkelsen, bjørn thomasen, henrik vosgerau and stuart watt in east and north-east greenland and greg barnes, stefan bernstein, gregers dam, søren hansen, christian knudsen, lotte m. larsen, asger k. pedersen and chris pulvertaft in west greenland. practical help over many years and analytical support by john boserup, ditte kiel-dühring, carsten guvad and h. p. nytoft are highly appreciated. the preparation of the actual web-gis inventory and the supporting manuscript was made with much help from ane d. asmussen, jette halskov, jonas del pin hamilton, tjerk heijboer, simun d. olsen, diogo rosa, lisbeth tougaard and willy l. weng. thanks to jolanta kus for providing unpublished information from bgr. results are from many different field projects by ggu/geus with additional funding from danish and greenlandic authorities. contributions came especially from the danish energy research program, the previous mineral resources administration for greenland in copenhagen, the previous bureau of minerals and petroleum in nuuk and most recently the department of foreign affairs and energy. support from industry in east and north-east greenland and on disko and nuussuaq has been beneficial. special appreciation is directed to niels henriksen, ‘oscar’, the driving force behind the modern geological mapping of greenland, leader of countless expeditions and a never-failing source of information and inspiration. additional information funding statement the web-based gis inventory and the present paper were based on funding from the department of foreign affairs and energy in nuuk and geus, respectively. author contributions fgc: field descriptions and completion of first draft of the paper. jabk: field descriptions, paper writing and provided analytical data. https://doi.org/10.34194/geusb.v47.6519 http://www.geusbulletin.org christiansen & bojesen-koefoed 2021: geus bulletin 47. 6519. https://doi.org/10.34194/geusb.v47.6519 17 of 20 www.geusbul let in.org competing interests the authors declare no competing interest. the authors do not accept any responsibility for the use of information provided here in future exploration. additional files all additional materials with detailed locality descriptions are found in the web-based gis inventory: https://doi.org/10.22008/fk2/jkyxjm. references abrams, m.a. 2005: significance of hydrocarbon seepage relative to petroleum generation and entrapment. marine and petroleum geology 22, 457–477. https://doi.org/10.1016/j.marpetgeo.2004.08.003 alexandre, p. & kyser, k.t. 2006: geochemistry of uraniferous bitumen in the southwest athabasca basin, saskatchewan, canada. economic geology 101, 1605–1610. https://doi.org/10.2113/gsecongeo.101.8.1605 andrews, s.d., decou, a., braham, b., kelly, s.r.a., robinson, p., morton, a., marshall, j.e.a. & hyden, f. 2020a: exhumed hydrocarbon traps on the north atlantic margin: stratigraphy, palaeontology, provenance and bitumen distribution, an integrated approach. basin research 32(5), 1223–1243. https://doi.org/10.1111/bre.12424 andrews, s.d., decou, a., braham, b., kelly, s.r.a., robinson, p., morton, a., marshall, j.e.a. & hyden, f. 2020b: exhumed hydrocarbon traps in east greenland: reply to christensen et al.’s comment on andrews et al. 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copper-lead-zinc mineralisation in the permo-triassic of central east greenland. bulletin grønlands geologiske undersøgelse 143, 42 pp. https://doi.org/10.34194/bullggu.v143.6685 thrane, k. & kalvig, p. 2019: graphite potential in greenland. geology and ore 32, 12 pp. van der stijl, f.w. & mosher, g.z. 1998: the citronen fjord massive sulphide deposit, peary land, north greenland: discovery, stratigraphy, mineralization and structural setting. geology of greenland survey bulletin 179, 40 pp. https://doi.org/10.34194/ggub.v179.6270 von guttenberg, r. & van der stijl, f. 1993: north greenland project 1992. report of work. unpublished company report. nanisivik mines ltd., platinova a/s. 129 pp. toronto. watt, w.s. 2019: stratigraphic lexicon for greenland. 327 pp. copenhagen: geological survey of denmark and greenland. https://doi. org/10.22008/geusbook/strat-lex-greenland watt, w.s. & wrang, p. 1984: migrated hydrocarbons in basalt on the south side of scoresby sund. rapport grønlands geologiske undersøgelse 120, 84–85. https://doi.org/10.34194/rapggu.v120.7863 zheng, g., xua, w., etiope, g., ma, x., liangd, s., fana, q., sajjada, w. & lia, y. 2018: hydrocarbon seeps in petroliferous basins in china: a first inventory. journal of asian earth sciences 151, 269–284. https://doi. org/10.1016/j.jseaes.2017.10.037 https://doi.org/10.34194/geusb.v47.6519 http://www.geusbulletin.org https://doi.org/10.34194/rapggu.v58.7361 https://doi.org/10.34194/bullggu.v143.6685 https://doi.org/10.34194/ggub.v179.6270 https://doi.org/10.22008/geusbook/strat-lex-greenland https://doi.org/10.22008/geusbook/strat-lex-greenland https://doi.org/10.34194/rapggu.v120.7863 https://doi.org/10.1016/j.jseaes.2017.10.037 https://doi.org/10.1016/j.jseaes.2017.10.037 inventory of onshore petroleum seeps and stains in greenland: a web-based gis model abstract 1 introduction 2 classifications and nomenclature 3 inventory 4 regional distribution of petroleum seeps and stains in greenland 5 north greenland 5.1 petroleum seeps and stains 5.2 gas 5.3 bitumen associated with mineral occurrences 6 east and north-east greenland 6.1 petroleum seeps and stains 6.2 gas 6.3 bitumen associated with mineral occurrences 7 central west greenland 7.1 petroleum seeps and stains 7.2 gas 7.3 bitumen associated with mineral occurrences 8 south and southern west greenland 9 offshore 10 conclusions and recommendations acknowledgements additional information funding statement author contributions competing interests additional files references fig. 6 stratigraphy of west greenland (nuussuaq basin) with simplified relations between petroleum s fig. 5 stratigraphy of east and north-east greenland with simplified relations between petroleum see fig. 4 stratigraphy of north greenland (franklinian basin) with simplified relations between petrole fig. 3 collection of field photos showing petroleum seepage and stains from greenland. all photos ar fig. 2 close up examples from the web-based gis model, in this case with a topographic map as backgr fig. 1 map of greenland showing distribution of petroleum steeps and stains. red triangles: petrole geological survey of denmark and greenland bulletin 35, 2016, 95-98 95© 2016 geus. geological survey of denmark and greenland bulletin 35, 95–98. open access: www.geus.dk/publications/bull in april 2015, the geological survey of denmark and greenland (geus) together with the ministry of mineral resources in greenland (mmr) made further progress in the development of web-based facilities to present and disseminate geoscientific information and data. this was presented in a new version of the greenland mineral resources portal. the portal now provides the users with access to a wealth of geological, geophysical and geochemical data – mostly free of charge (fig. 1). the primary goal of the portal is to facilitate data searches for exploration companies, but the benefits from the easy access to geoscience data are also open to academic researchers. it is the plan to add several new functionalities to the portal in 2016, and interested users are invited to visit the portal at www. greenmin.gl on a regular basis to follow progress. geoscience data – from archive to internet geus, and its forerunner in greenland (ggu), had as its core activity the production of geological maps. in the early days, observations and data were documented in field notes and on paper maps, but since the 1980s more and more field data have been recorded digitally. old archive data are still valuable because in greenland data acquisition is expensive. therefore geus has worked on digitising and securing historic data and integrating them into modern databases with access through the internet. geus has placed geological data from greenland on the internet since 2005, when the greenland mineral occurrence map database (gmom) was first opened to the public (thorning et al. 2004). initially, the gmom portal expanded geographically year by year, providing access to mineral occurrence data extracted by geus from released company reports. in 2011, a new, more ambitious programme was launched as a collaborative project between geus and the bureau of minerals and petroleum in greenland (bmp), the forerunner of mmr. the objective was to compile data and maps from the whole of greenland into a modern web-platform – the greenland mineral resources portal – with access to large amounts of the greenland mineral resources portal – another step forward mikael pedersen, martin hansen, bjørn h. heincke and leif thorning fig. 1. the interactive map of the greenland mineral resources portal showing the mineral occurrence layer. 9696 geoscientific metadata, including full texts of reports made for companies from the dodex database (riisager et al. 2011). the first version of this portal was released in the beginning of 2012, and mainly provided information about the existence of data in different areas, but only to a limited degree the data themselves. however, the need for easy data access for industry was obvious and increasingly the actual data from the large range of geus geological databases were added. i n 2015 new content and functions were launched to give the users access to download or purchase geochemical and geophysical data and a range of other data types such as geological maps and sample data. high-quality data geological data are essential for development in the mining sector. factors such as data quality and access to databases are therefore important parameters when rating a country’s attraction to investors as expressed by, for example, the annual survey of mining companies undertaken by the fraser institute (jackson & green 2016). in its capacity as a national geological data and knowledge centre, it has for many years been geus’ fundamental task to develop databases securing valuable data assets. high-quality databases, however, rely not only on skilled developers, but also very much on geoscientific personnel that can continuously ensure the quality of the content. all data have been quality controlled in the central, authoritative databases before being stored and made available in the greenland mineral resources portal. a few of the data types are described below. geophysical data. geophysical data are in great demand for mineral exploration and the integration of such data into the portal has therefore been a high priority. during the past two decades, geus and bmp/mmr have jointly acquired airborne geophysical data in greenland within the aem greenland and aeromag projects (rasmussen et al. 2013). geophysical data are regularly reported by exploration companies to the greenland authorities as part of their licensing conditions and these are released to the public after a confidentiality period of five years. prior to 2011, these surveys were handled file by file, but an important achievement of the portal project has been to develop a relational database for the geophysical metadata and data files to make them identifiable for purchase or download via the portal. this work is ongoing, and at present a total of 35 geophysical surveys are available. geochemical data. stream sediment samples have been collected in greenland since 1971 and geochemical mapping by means of systematically collected and analysed stream sediment samples has been undertaken by ggu/ geus since 1975. a geochemical atlas of west and south greenland was published in 2001 (steenfelt 2001a) based on 7122 samples analysed for 43 chemical elements. a major task in the preparation of the atlas was the elimination of bias in the analytical data as a prerequisite for the production of the element distribution maps (steenfelt 1999, 2001b). in 2011, a similar exercise was performed on five elements determined in 2644 stream sediment samples from north greenland as part of an assessment of the zinc fig. 2. northernmost greenland as it appears in the interactive geological map of greenland on a scale of 1:500 000. 97 potential (thrane et al. 2011). all geochemical data from these two projects are available in the greenland mineral resources portal – via the interactive map interface and as free download from the associated webshop. an important part of the ongoing project is the quality control of all geochemical analyses from more than 15 000 raw (noncalibrated) stream and scree sediment samples as well as from soil samples and heavy mineral concentrates – all of which are part of the portal. geological maps. geological mapping of greenland has led to a series of published maps at standard scales. the maps were traditionally produced using engraving techniques, but in 1998 geus turned to digital production of paper maps. in recent years the old paper maps have been digitised and made into seamless, geographically referenced products with homogeneous legends, suitable for web applications. a seamless 1:100 000 scale map with a homogenised legend covering 10 map sheets in southern west and south-west greenland was finalised in 2010 (keulen et al. 2010), and in 2013 geus launched a 1:500 000 scale compilation covering the whole of the ice-free area of greenland (pedersen et al. 2013; fig. 2). both map layers form important base maps in the greenland mineral resources portal. the gis data from the 1:500 000 scale map are available for free download from the webshop. targeting the end users the development of digitised facilities has been going on for many years during which valuable experience has been gathered. the development has always been demanddriven, but even though exploration companies are a quite well-defined user group, various requirements need to be addressed. the focus is not only to provide access to as much reliable data as possible, but also to create online facilities for screening and evaluating the data. therefore, the development of the portal is based on cases that describe scenarios for a specific user (e.g. a small-scale miner) expecting a certain result. a couple of scenarios are given in the following to demonstrate the flexibility and potential of the greenland mineral resources portal. scenario 1: discovery and download of geophysical data in a given area. many exploration geologists are interested in geophysical data from a specific area. in the interactive map of the greenland mineral resources portal it is easy to zoom in on an area and see the available aeromagnetic data by switching on the total magnetic intensity layer (fig. 3a). by adding the geophysical surveys for sale layer to the map, information about the individual geophysical surveys available in the area, including metadata and reports, can be inspected (fig. 3b). the user buys data by adding the relevant surveys to the shopping basket, after which they can be checked out from the webshop (fig. 3c). many geophysical datasets are available free of charge, and others can be paid by credit card and subsequently downloaded. scenario 2: search for stream sediment samples with high gold content. in the interactive map users can search through each of the geochemical datasets for samples with an element content larger than a value defined by the user. this makes it possible for the user to easily assess areas of interest a b c fig. 3. a: the total magnetic intensity layer shows where airborne magnetic data are available. b: the layer of geophysical surveys for sale provides information on the individual geophysical surveys. intuitive interaction between search results and the map makes it easy to see what geographical areas are part of the individual surveys. based on metadata and the access to survey reports the user can assess the surveys before purchase or download. c: individual surveys can be added to the online shopping basket and checked out through the webshop. 9898 without advanced gis software. in fig. 4 the result of a search for samples with more than 25 ppb gold is shown for an area in south-west greenland. the locations of stream sediment samples that fulfil that criterion are displayed on a map, and more information about the samples can be seen by clicking the red dots. more advanced users can download all of the geochemical datasets (or geographic subsets) free of charge from the webshop for further analyses using their own software. looking ahead the development of the greenland mineral resources portal is ongoing. more and more data types are added, and the existing datasets are continuously supplemented. in 2016, it is the plan to add a geochronological database and a drill-core directory and future plans include integration of satellite imagery, additional aerial photographic collections, more geochemical and geophysical data and digital elevation models. the user-friendly functions of the portal interface will be further improved and refined along the way, and users are always encouraged to give feed-back to geus and the ministry of mineral resources for continuous improvements. acknowledgements the ministry of mineral resources under the government of greenland is thanked for financial support for the development of the greenland mineral resources portal. the authors also wish to thank the team from the department of petrology and economic geology as well as the geological datacentre at geus, who have actively worked to make the greenland mineral resources portal a success. references jackson, t. & green, k.p. 2016: fraser institute annual survey of mining companies 2015, 86 pp. vancouver: fraser institute. keulen, n.t., kokfelt, t.f. & scherstén, a. 2010: notes on the common legend to the 1:100 000 digital geological map of southern west and south-west greenland, 61º 30´–64°n. danmarks og grønlands geologiske undersøgelse rapport 2010/119, 41 pp. pedersen, m., weng, w.l., keulen, n.t. & kokfelt, t.f. 2013: a new seamless digital 1:500 000 geological map of greenland. geological survey of denmark and greenland bulletin 28, 65–68. rasmussen, t.m., thorning, l., riisager, p. & tukiainen, t. 2013: airborne geophysical data from greenland. geology and ore, exploration and mining in greenland 22, 12 pp. riisager, p., pedersen, m., jørgensen, m.s., schjøth, f. & thorning, l. 2011: dodex – geoscience documents and data for exploration in greenland. geological survey of denmark and greenland bulletin 23, 77–80. steenfelt, a. 1999: compilation of data sets for a geochemical atlas of west and south greenland based on stream sediment surveys 1977 to1997. danmarks og grønlands geologiske undersøgelse rapport 1999/41, 33 pp. steenfelt, a. 2001a: geochemical atlas of greenland – west and south greenland. danmarks og grønlands geologiske undersøgelse rapport 2001/46, 39 pp. steenfelt, a. 2001b: calibration of stream sediment data from west and south greenland. a supplement to geus report 1999/41. danmarks og grønlands geologiske undersøgelse rapport 2001/47, 43 pp. thorning, l., christensen, l., schjøth, f. & stendal, h. 2004: greenland mineral occurrence map. status report for the development of a prototype for the internet, january 2004. danmarks og grønlands geologiske undersøgelse rapport 2004/28, 52 pp. thrane, k., steenfelt, a. & kalvig, p. 2011: zinc potential in north greenland. danmarks og grønlands geologiske undersøgelse rapport 2011/143, 64 pp. fig. 4. a search for samples with more than 25 ppb gold in the calibrated stream sediment dataset from south-west greenland. all analytical results for a given sample can be seen on the screen, and the entire (or geographic subsets) of the dataset can be downloaded from the webshop. authors’ address geological survey of denmark and greenland), øster voldgade 10, dk-1350 copenhagen k, denmark, e-mail: mp@geus.dk research article | short dahl et al. 2023: geus bulletin 53. 8357. https://doi.org/10.34194/geusb.v53.8357 1 of 7 neural network predictions of drawdown from groundwater abstraction in the egebjerg catchment, denmark mathias busk dahl*1,2 , troels norvin vilhelmsen2 , trine enemark3 , thomas mejer hansen1 1department of geoscience, aarhus university, aarhus, denmark; 2niras, aarhus, denmark; 3department of geosciences and natural resource management, university of copenhagen, copenhagen, denmark abstract results from numerical simulations play a vital role in the decision process of everyday groundwater management. however, these simulations can be time-consuming for large-scale investigations, and it can be necessary to apply approximate methods instead. this study investigates the abilities of a neural network to replicate simulated drawdown from groundwater abstraction in a numerical groundwater model of the egebjerg catchment, denmark. we follow a generalised methodology that uses the information within the deterministic numerical model to create a training set for the neural network to learn from and extend the method to work in a 3d danish groundwater model case. we compare the abilities of the trained neural network with the results of conventional computations in terms of speed and accuracy and argue that this approach has the potential to improve decision support for decision-makers within groundwater management. *correspondence: dahl@geo.au.dk received: 20 jun 2023 revised: 18 aug 2023 accepted: 22 sep 2023 published: 10 nov 2023 keywords: decision support, groundwater modelling, machine learning, probabilistic neural network, resource management abbreviations drn: drain package ghb: general head boundary package gelu: gaussian error linear unit rch: recharge package riv: river package wel: well package geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: hyojin kim (geus, denmark) reviewed by: robin thibaut (ghent university, belgium) and two anonymous reviewers funding: see page 6 competing interests: see page 6 additional files: see page 6 introduction groundwater models informed by data on geology and hydrological properties of the subsurface are important tools in groundwater management (gorelick 1983; pisinaras et al. 2007; hadded et al. 2013). these models simulate groundwater flow in the subsurface and are used to investigate the environmental effects of external interventions on the groundwater system, such as the establishment of new abstraction wells. dahl et al. (2023) provide an approach for training a neural network to replicate the results of drawdown from an abstraction well as simulated in a modflow model (harbaugh 2005) with a probabilistic output. the network is trained to model the mapping between a few influential model attributes and simulated drawdown. once trained, the network carries out this mapping at a speed exceeding 100 times that of conventional modflow simulations. dahl et al. (2023) make use of a synthetic groundwater model and limit the drawdown predictions to the layer of pumping. here, we extend the method presented in dahl et al. (2023) to allow abstraction from multiple suitable model layers and predict drawdown in a full 3d real field-based groundwater model of the egebjerg catchment, east jylland, denmark (approx. 55.9748°n, 9.8129°e to 55.8581°n, 9.9645°e). predictions from the neural network are compared with modflow simulation results, and we discuss the generalisation potential of the tested method. materials and methods the methodology of this paper is based on that presented by dahl et al. (2023). our aim is to generalise and test the method on a danish groundwater model case and explore its implications for groundwater management. https://doi.org/10.34194/geusb.v53.8357 https://orcid.org/0000-0003-4816-0241 https://orcid.org/0000-0002-6399-8563 https://orcid.org/0000-0002-3881-2739 https://orcid.org/0000-0003-4529-0112 mailto:dahl@geo.au.dk https://creativecommons.org/licenses/by/4.0/deed.ast dahl et al. 2023: geus bulletin 53. 8357. https://doi.org/10.34194/geusb.v53.8357 2 of 7 www.geusbul let in.org the core idea is to train a neural network to predict drawdown from groundwater abstraction using results and features from a numerical groundwater model to significantly reduce computation times. whilst the method can generally be applied to most groundwater models, a few site-specific decisions need to be made. our objective is to make minimal changes to the method to accommodate the new groundwater model. we also aim to expand the method’s capabilities from 2d to a full 3d model by training the neural network with simulated drawdown data from all model layers, which come from wells situated in various aquifers. all the work is performed on a computer with an 13th gen intel core i9 3.00 ghz processor and 128 gb ram. egebjerg groundwater model for the egebjerg catchment, a groundwater flow model has been developed using modflow 6 (langevin et al. 2017, 2019). this model employs a grid size of 144 × 121 with a 100 × 100 m discretisation across its 14 layers of varying thicknesses (enemark et al. 2022). boundary conditions are modelled using the general head boundary package (ghb) for lakes and coastal areas, the recharge package (rch) for the simulation of groundwater recharge in the top active cells, the drain package (drn) for simulating stream flow (drn-riv in fig. 1a) and drainage in all top active cells and the well package (wel) for simulating groundwater abstraction in active wells. topography and the location of the boundary conditions in the model are visualised in fig. 1a. the geological layers alternate between sand and clay deposits with a thick chalk aquifer beneath. the downwards extent of the chalk layer is unknown but is assumed to be –550 m below sea level. to simulate layers that pinch out, the vertical pass-through option in modflow 6 is used. this option is applied for cells with a thickness of <0.5 m where flow is distributed downwards to the subsequent active cell. data-set construction the data set for training the neural network consists of sets of target data y and input features stored in a vector x. features are related to targets with an unknown function f, such that: y = f (x), (1) where f is determined through training of the neural network (gardner & dorling 1998). here, the target data y in eq. (1) is the simulated drawdown caused by groundwater abstraction in the model. the training data set is created by conducting 1000 simulations. in each simulation, a new well is introduced into the groundwater model. the well’s location is randomly chosen from one of the layers containing water and having a thickness >10 m. additionally, the pumping rate for the well is drawn randomly from a uniform range spanning 100–5000 m3·day–1. a single modflow simulation takes 3.2 ± 0.1 s to run. the change in drawdown in each cell caused by the abstraction is saved for the target data. dahl et al. (2023) proposed to use a subset of the full information in the model for the input feature vector x. the subset contains 12 influencing features that are generally available in most groundwater models. we extend the method from only observing drawdown from abstraction in a single layer to the full 3d model, where a well can be placed at different depths, and changes in drawdown are predicted for all layers. the features from the egebjerg groundwater model used for training include the hydraulic head before pumping, distance to well (fig. 1b), travel time to well (fig. 1c), distance to head boundary (ghb package; fig. 1d), distance to stream (drn package; fig. 1e), pumping rate, hydraulic conductivity and the logarithmic hydraulic conductivity in the current cell, well location (row, column and layer) and the layer of the current cell. the travel time feature does not represent a specific physical measure but is a proxy of water’s ability to travel to a certain location depending on the distance and flow resistance along the travel route (hydraulic conductivity). the fast-marching algorithm is used to compute these travel times in a 2d plane of the active cells just below topography, as implemented in the scikit-fmm python extension module (furtney 2021) and applied in other studies for training machine learning models with simulated data (thibaut et al. 2021). neural network setup the structure of the neural network consists of an input layer of size 12 (the number of input features), fully connected to three hidden layers, each with 75 neurons as in dahl et al. (2023). the output layer has two neurons for estimating the mean and standard deviation of a normal distribution representing the drawdown. we apply the gaussian error linear unit (gelu) activation function (hendrycks & gimpel 2016) between the input and hidden layers and a linear activation function between the last hidden layer and the output layer. the adam algorithm (kingma & ba 2014) is used to minimise the loss function, which is defined as the negative log-likelihood of 1d gaussian distribution. this allows us to interpret the output of the neural network as a gaussian distribution as describing the drawdown (dahl et al. 2023). we use the python libraries tensorflow and tensorflow probability (abadi et al. 2016; dillon et al. 2017) for the construction and training of the neural network. https://doi.org/10.34194/geusb.v53.8357 http://www.geusbulletin.org dahl et al. 2023: geus bulletin 53. 8357. https://doi.org/10.34194/geusb.v53.8357 3 of 7 www.geusbul let in.org the holding input features of the constructed data set and simulated modflow responses are split 80/20 into a training data set and a validation data set and standardised using a standard scaler. the validation set is withheld from training and used to validate the ongoing training phase to prevent overfitting. during training, if loss improves for the training data but not for the validation data, the neural network is likely overfitting to the training data. we train the network for 1000 epochs and observe a stagnation in loss improvement without encountering overfitting. additionally, an independent third test data set is constructed from 80 new modflow simulations, separate from the validation data set for performance evaluation. results the performances of the trained neural network on the validation and test data are shown in fig. 2a–b as normalised density plots of modflow simulations (observed drawdown) against mean predictions (predicted drawdown) with a logarithmic colour scale. the sum of all values in the figure adds up to 1. the dashed identity line represents the ideal agreement between modflow simulations and predictions of the neural network. in both data sets, we observe the highest concentrations exactly on the identity line for a drawdown from 0 m and up to c. 3 m. here lies the highest density region containing 95% of the data marked with the white contour line (fig 2a–b). in this interval, the data density rapidly decreases as we move away from the line in both directions. for higher values, we observe a shift in the validation data density plot away from the identity line, and the spread increases. figure 2c shows histograms depicting the differences between true modflow results and predictions within the two data sets, confirming a small error of less than 0.2 m for most differences. figure 2d displays histograms depicting data set z scores, which quantify the deviation of modflow results xmod from the mean prediction μnn in terms of standard deviations, σnn where, (2) the root mean squared errors between observed and predicted values in fig. 2a–b are rmsevali  = 0.19 m and rmsetest = 0.17 m. the low errors are in good agreement with the observed high density of observations near the identity line and the distribution in fig. 2c. we test the neural network at two different locations outside the original training and validation data sets and compare predictions to modflow results (fig. 3). in case 1 (fig. 3a), the well is placed in layer 6 at row 55 and column 99 with a pumping rate of 798 m3·day–1. the drawdown from the modflow simulation shown in fig. 3a is compared to the predicted mean and standard deviation from the network (fig. 3b–c). the difference between modflow and the predicted mean of 0 20 40 60 80 100 120 column 0 20 40 60 80 100 120 140 r ow a boundary conditions and topography 0 50 100 150 topography [m] drn-riv ghb well 0 50 100 column 0 25 50 75 100 125 r ow distance to well well well wellwell b d e c 0 50 100 column 0 25 50 75 100 125 r ow travel time 0 50 100 column 0 25 50 75 100 125 r ow distance to ghb 0 50 100 column 0 25 50 75 100 125 r ow distance to stream 0 2000 4000 6000 8000 10000 d istance [m ] 50000 100000 150000 200000 250000 300000 350000 400000 tim e [d ays] 0 1000 2000 3000 4000 d istance [m ] 0 1000 2000 3000 4000 d istance [m ] input features fig. 1 a visualisation of the egebjerg model boundary conditions and some of the relevant features for training. a: streams (blue; drn-riv), lakes and fjords (green et al. 2011) and wells (red; well) are modelled as boundary conditions shown on top of a topographic map. b–e: examples of input features to the neural network. drn-riv: drain package. ghb: general head boundary package. https://doi.org/10.34194/geusb.v53.8357 http://www.geusbulletin.org dahl et al. 2023: geus bulletin 53. 8357. https://doi.org/10.34194/geusb.v53.8357 4 of 7 www.geusbul let in.org the network is visualised in fig. 3d. the network replicates the modflow results well with low differences between the two in most parts of the layer. we observe some areas with larger differences that seem to correlate with locations of boundary conditions such as simulated rivers and other wells. also, positive values (in fig. 3d) are observed near the well, indicating that the neural network slightly underestimates the drawdown in this area. these difficult areas are also visible in the standard deviation map of the neural network where locations of other wells are clearly observed. in case 2 (fig. 3e), the well is placed in layer 6 at row 85 and column 47 with a pumping rate of 849 m3·day–1. again, modflow results are compared to the neural network (fig. 3f–g) and subtracted to show the difference (fig. 3h). the network models the overall extent of the change in layer, though the discrepancy between modflow and network predictions is larger compared to case 1. high differences are especially apparent at the nearby boundary conditions. again, this difference between results is correlated to the standard deviation map in fig. 3g, where larger standard deviations are observed. 0 1 2 3 4 5 6 7 8 9 predicted drawdown [m] 0 1 2 3 4 5 6 7 8 9 o bs er ve d dr aw do w n [m ] a validation data per formance 95% contour line identity l ine 0 1 2 3 4 5 6 7 8 9 predicted drawdown [m] 0 1 2 3 4 5 6 7 8 9 o bs er ve d dr aw do w n [m ] b test data per formance 95% contour line identity l ine true predicted [m] 0 5 10 15 20 25 d en si ty c difference between true and predicted head test data validation data z-score 0.00 0.05 0.10 0.15 0.20 0.25 0.30 0.35 0.40 d en si ty d z-scores of test and validation data test data validation data log density log density 10 – 6 10 – 5 10 – 4 10 –3 10 –2 10 –1 10 – 6 10 – 5 10 – 4 10 –3 10 –2 10 –1 –0.6 –0.4 0.0–0.2 0.2 0.4 0.6 –10.0 –7.5 –2.5–5.0 0.0 2.5 5.0 7.5 10.0 fig. 2 the performances of the trained neural network on the validation data set and an independent test data set are shown as density plots a and b. the x-axis represents the predicted mean values of the drawdown, whilst the y-axis represents the observed modflow drawdown. each square in the figures covers several observations visualised with a normalised logarithmic colour scale. the identity line (blue dashed line) shows the one-toone agreement between modflow and the network. the white contour line holds 95% of the data within. c shows the difference between the true modflow values and predicted values for both data sets as a probability density histogram. d depicts the distribution of z scores for the validation and test data sets. brown shading: overlap between test data and validation data sets. https://doi.org/10.34194/geusb.v53.8357 http://www.geusbulletin.org dahl et al. 2023: geus bulletin 53. 8357. https://doi.org/10.34194/geusb.v53.8357 5 of 7 www.geusbul let in.org predicting the full model drawdown with the network takes 0.60 ± 0.01 s, related to running the neural network and input-feature creation. the run-time is 3.2 ± 0.1 s for modflow simulations. with the pumping rate as an input feature, it is possible to perform tests of pumping series where pumping rates are gradually increased, and drawdown is predicted using the neural network. figure 4 shows a drawdown in multiple layers for this test where two different pumping rates are used with the same well setup as in case 1 (fig. 3a–d). here, all input features are determined once, and only the pumping rate feature is changed for all following network predictions. this reduces the run-time to 0.29 ± 0.04 s per simulation. discussion we have trained a neural network to replicate the results of simulated drawdown using modflow due to groundwater abstraction in the egebjerg groundwater 0 20 40 60 80 100 120 140 r ow well pos. layer: 6 row: 55 col.: 99 a modflow b neural network mean c neural network std. d difference 0 50 100 column 0 20 40 60 80 100 120 140 r ow well pos. layer: 6 row: 85 col.: 47 e modflow 0 50 100 column f neural network mean 0 50 100 column g neural network std. 0 50 100 column h difference 0.0 0.2 0.4 0.6 0.8 1.0 d raw dow n [m ] –0.2 –0.1 0.0 0.1 0.2 d ifference [m ] 0.0 0.2 0.4 0.6 0.8 1.0 d raw dow n [m ] –0.2 –0.1 0.0 0.1 0.2 d ifference [m ] fig. 3 results and comparisons between modflow and the neural network in two test cases. a–d: case 1. a new well is simulated in layer 6 at row 55 and column 99. the drawdown of the layer above the abstraction layer is visualised for modflow (a), the neural network mean and one standard deviation (b–c) and the difference between modflow and network mean (d). e–h: case 2. the well is moved to layer 6 at row 85 and column 47, and drawdown is estimated again as in a–d. row 0 20 40 60 80100120140 column 0 20 40 60 80 100 120 layer 2 layer 5 layer 14 0.0 0.2 0.4 0.6 0.8 1.0 d raw dow n [m ] row 0 20 40 60 80100120140 column 0 20 40 60 80 100 120 layer 2 layer 5 layer 14 pumping rate = 800 m3/day 0.0 0.2 0.4 0.6 0.8 1.0 d raw dow n [m ] pumping rate = 200 m3/day fig. 4 drawdown in multiple layers predicted with the neural network with the well setup from case 1. left: the pumping rate is set to 200 m3·day–1. right: the pumping rate is increased to 800 m3·day–1. https://doi.org/10.34194/geusb.v53.8357 http://www.geusbulletin.org dahl et al. 2023: geus bulletin 53. 8357. https://doi.org/10.34194/geusb.v53.8357 6 of 7 www.geusbul let in.org model, taking the same approach as dahl et al. (2023) with only a few changes to input features. this demonstrates that the methodology of dahl et al. (2023) can be generalised, with only minor modifications when applied to new areas. to do this, we extend the method to predict drawdown in all layers and allow the well to be placed in multiple, suitable layers. as a performance test of the neural network on a validation set and an independent test set, results of drawdown from multiple well locations in different layers are presented in fig.  2. the outcomes reveal a strong agreement between the predictions of the neural network and the results obtained from modflow (fig. 2a–c), and that a substantial amount of the error can be described with 2 standard deviations (fig. 2d). the neural network shows a tendency to underestimate higher values of drawdown in fig. 2a. this is also the case in dahl et al. (2023) and is likely a consequence of a few high-value observations in the training data compared to the large number of low values available. this could be solved by sampling differently from the training data to favour underrepresented data (johnson & khoshgoftaar 2019) or by error-correcting biased predictions (belitz & stackelberg 2021). the case analyses from fig. 3 show that the network in general replicates the modflow results well. some difficulties near simulated boundary conditions have a large effect on the predicted mean values. however, these areas also show an increase in the standard deviation maps making them easier to spot and possibly remove or ignore. the standard deviation could potentially act as a verifier of the mean estimate and identify network shortcomings. the accuracy of the network confirms that the applied approach generalises to a danish case and likely other danish groundwater models. the neural network predictions are obtained five times faster than modflow results in the full model. this modest speed-up can be attributed to the efficient run-time of the egebjerg groundwater model combined with the computation of input features but could potentially be much higher in a model with higher resolution. the neural network, however, provides inherent flexibility, allowing for predictions within specific subareas of the model rather than necessitating computations for the entire model each time, thereby reducing the computational load. furthermore, for the pumping analysis (fig. 4), all features, except for the pumping rate, need to be computed only once. this results in subsequent network simulations being 11 times more efficient than those using modflow. such an analysis for a subarea of the model could greatly reduce the computational time compared to modflow. the proven speed-up and replication accuracy of the network make it a great addition to decision support tools, for example, during initial screening investigations. conclusions in this study, we have trained a neural network to predict drawdown from groundwater abstraction to test the abilities of the network and the applied approach to generalise to a danish catchment. we extend the approach to the full 3d model to increase user flexibility and general applicability with only a few modifications to input features and network setup. we find a good agreement between modflow results and network predictions and show that areas with high disagreement correlate well with larger network standard deviations. the neural network offers a 5-time speed-up compared to modflow runtime for a single simulation and 11-time speed-up in pumping analysis tests with multiple model runs where input features are only calculated once. furthermore, the network is flexible and can be limited to predict changes in subareas of the model reducing the number of computations required. we conclude that the applied approach generalises well to the danish groundwater model, and that the trained network has potential as a decision support tool. acknowledgements the authors would like to acknowledge innovation fund denmark for funding this project. the authors would also like to thank robin thibaut and two anonymous reviewers for a constructive review process that improved the quality of this paper. and a special thanks to handling editor, hyojin kim. additional information funding statement this work was funded by the innovation fund denmark, project 9065-00212b. author contributions md: conceptualisation, code development, investigation, methodology, validation, visualisation, writing – original draft, writing – review & editing. tv: conceptualisation, methodology, validation, supervision, writing – review & editing. te: code development, model creation, writing – review & editing. th: conceptualisation, methodology, validation, supervision, funding acquisition, writing – review & editing. competing interests tv and md were employed by niras. the remaining authors declare that the research was conducted in the absence of any commercial or financial relationship that could be construed as a potential conflict of interest. additional files data and code available at https://github.com/mathiasbusk/ hydrosim_egebjerg. references abadi, m. et al. 2016: tensorflow: large-scale machine learning on heterogeneous distributed systems. arxiv preprint arxiv:1603.04467. https://doi.org/10.48550/arxiv.1603.04467 https://doi.org/10.34194/geusb.v53.8357 http://www.geusbulletin.org https://github.com/mathiasbusk/hydrosim_egebjerg https://github.com/mathiasbusk/hydrosim_egebjerg https://doi.org/10.48550/arxiv.1603.04467 dahl et al. 2023: geus bulletin 53. 8357. https://doi.org/10.34194/geusb.v53.8357 7 of 7 www.geusbul let in.org belitz, k. & stackelberg, p.e. 2021: evaluation of six methods for correcting bias in estimates from ensemble tree machine learning regression models. environmental modelling & software 139, 105006. https://doi. 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https://doi.org/10.1016/j.jhydrol.2011.05.002 https://doi.org/10.1007/s11269-013-0266-7 https://doi.org/10.3133/tm6a16 https://doi.org/10.3133/tm6a16 https://doi.org/10.48550/arxiv.1606.08415 https://doi.org/10.48550/arxiv.1606.08415 https://doi.org/10.48550/arxiv.1412.6980 https://doi.org/10.48550/arxiv.1412.6980 https://doi.org/10.1186/s40537-019-0192-5 https://doi.org/10.1186/s40537-019-0192-5 http://pubs.er.usgs.gov/publication/tm6a55 https://doi.org/10.5066/f76q1vqv https://doi.org/10.5066/f76q1vqv https://doi.org/10.1007/s10666-006-9040-z https://doi.org/10.1016/j.jhydrol.2021.126903 https://doi.org/10.1016/j.jhydrol.2021.126903 neural network predictions of drawdown from groundwater abstraction in the egebjerg catchment, denma 1. introduction 2. materials and methods 2.1 egebjerg groundwater model 2.2 data-set construction 2.3 neural network setup 3. results 4. discussion 5. conclusions acknowledgements additional information funding statement author contributions competing interests additional files references figures fig. 1 a visualisation of the egebjerg model boundary conditions and some of the relevant features fig. 2 the performances of the trained neural network on the validation data set and an independent fig. 3 results and comparisons between modflow and the neural network in two test cases. a-d: case 1 fig. 4 drawdown in multiple layers predicted with the neural network with the well setup from case 1 geological survey of denmark and greenland bulletin 17, 2009, 41-44 groundwater mapping in denmark has high priority. it was initiated in the 1990s when the pressure on groundwater resources increased due to urban development and pollution from industrial and agricultural sources. in some areas, the groundwater mapping included survey drillings, modelling based on existing knowledge and geophysical mapping with newly developed methods that made area coverage on a large scale possible. the groundwater mapping that included development of new geophysical methods showed promising results, and led to an ambitious plan to significantly intensify the hydrogeological mapping in order to improve the protection of the danish groundwater resources. in 1999 the danish government initiated the national groundwater mapping programme with the objective to obtain a detailed description of the aquifers with respect to localisation, extension, distribution and interconnection as well as their vulnerability to pollution (thomsen et al. 2004). this mapping programme covers around 40% of the area of den mark designated as particularly valuable water abstraction areas. water consumers fi nance the mapping programme by paying 0.04 € per cubic metre of consumed water. at the end of the programme in 2015, the total cost is estimated to be about 250 000 000 € with a significant part spent on geophysical mapping. the mapping programme is administered by seven local offices under the ministry of environment, but most of the practical work is carried out by private consulting companies, and involves the use of geophysical survey methods, survey drillings, well logging, water sampling and hydrological mapping, as well as geological and groundwater modelling. in major parts of the particularly valuable water abstraction areas, it is important to obtain spatially dense geophysical data covering large continuous areas. geophysical methods used in the hydrogeological mapping the choice of geophysical methods depends on the geological setting of the aquifers. those of interest for drinking water are primarily found within the upper 250 m of the subsurface. the aquifers can be grouped into three main types. in © geus, 2009. geological survey of denmark and greenland bulletin 17, 41–44. available at: www.geus.dk/publications/bull 41 fig. 1. areal extent of data collected by the end of 2008. a: areas with tem and skytem soundings, b: areas with paces profiles, c: cves profiles. d: seismic profiles. geophysical methods and data administration in danish groundwater mapping ingelise møller, verner h. søndergaard and flemming jørgensen c d jylland sjælland fyn a b 50 km fig. 3 rosa_2008:rosa-2008 01/07/09 15:48 side 41 the western part of denmark, extensive quaternary and prequaternary sand deposits dominate. in the central part, the most important groundwater resources are located in qua ternary sand deposits often found in quaternary valley structures deeply eroded into palaeogene clay deposits. in the northern and eastern parts of the country, most of the important aquifers are found in upper cretaceous and danian limestone. the most important geophysical methods are electrical and electromagnetic methods, combined with reflection seismic profiling and borehole logging at selected localities. differences in electrical properties between sandy aquifers and clay sediments favour the use of the electrical and electromagnetic methods (sørensen et al. 2005), but the ability of seismic methods to reveal detailed internal structures within the aquifers is also important. the most commonly used geophysical method in the groundwater mapping programme is the airborne transient electromagnetic method, skytem (sørensen & auken 2004), which is one of the new methods that has been developed to improve and optimise groundwater mapping. the first skytem groundwater mapping project was carried out in 2003. since then the skytem method has been developed further and has proved faster and more powerful than the groundbased, single-site transient electromagnetic method, tem, which was previously widely used. the skytem method is used for mapping to a maximum depth of 250–300 m. nu me rous buried valleys have been mapped in denmark by the tem method, in particular in the central parts of the country, where highly impermeable and low-resistive palaeogene clay layers form the lower boundaries of the aquifers and the valleys are easily detected. at the end of 2008, tem and skytem data cover an area of more than 11 000 km2 (fig. 1a), which is about one quarter of the area of denmark. electrical methods are used for near-surface mapping purposes. the pulled array continuous electrical sounding meth od (paces; sørensen 1996) has been extensively used to map layers within the upper 20–30 m. this method works well in combination with tem measurements, and the combined methods provide data from the surface down to 200–300 m. this combination of methods has mainly been used in eastern jylland and on fyn. a total of around 9000 effective line kilometres of paces data have been collected, corresponding to a coverage of more than 3000 km2 (fig. 1b). the continuous vertical electrical sounding method (cves; e.g. dahlin 1996) is used in areas where it is unne cessary to map deeper layers, and where the subsurface resistivity values are too high for the skytem method. about 4000 line kilometres have been collected, mainly in the central part of jylland and on the eastern part of sjælland (fig. 1c). the reflection seismic method is also of great value as a geophysical groundwater mapping tool, particularly following the development of a land-streamer and a new vibroseismic system (e.g. vangkilde-pedersen et al. 2006). although the reflection seismic method is expensive, it can be successfully combined with skytem measurements, and the decision about where to acquire seismic data can be based on the skytem results (jørgensen et al. 2003). successful mapping of the outline of buried valleys and their internal structures has been based on the interpretations of seismic profiles; skytem data do not allow such interpretations. the reflection seismic method has also been used successfully to map palaeogene and neogene sediments in the western part of denmark (rasmussen et al. 2007), where thick and extensive layers of sandy deposits constituting important aquifers are bounded by thinner layers of clayey deposits, and to map faults in danian and cretaceous limestone in the eastern part of denmark. around 1400 km of seismic lines have been collected, particularly in the western and central parts of jylland (fig. 1d). borehole logs are crucial for the geological and hydrological interpretation of boreholes. it is now common practise to log boreholes following survey drilling, and older water supply wells have also been logged. particularly in areas with chalk and limestone or neogene groundwater reservoirs log stratigraphy has provided valuable information. about 1500 boreholes have been logged. administration of the geophysical data the groundwater mapping programme is split up into many smaller areas to ease the administrative handling and to be able to meet priority criteria. careful and standardised treatment of data is required to ensure that the resulting ‘patchwork’ is of high and uniform quality and has no visible seams. therefore, standards and guidelines are worked out for geophysical data acquisition, calibration of instruments, data processing, interpretation (e.g. hydrogeophysics group 2007a) and geological modelling (jørgensen et al. 2008). without a predefined system of archiving the geophysical data and modelling results, the data logistics of the groundwater mapping programme would be overwhelming. the national geophysical relation database (gerda; http:// gerda.geus.dk) hosted at the geological survey of denmark and greenland (geus), is used for archiving these geophysical data. the development of the database began more than ten years ago. the database contains geophysical data of various types such as wenner profiles, schlumberger soundings, pulled array continuous electrical soundings, continuous vertical electrical soundings and induced polarisation, transient electromagnetic data including the airborne skytem data, 42 rosa_2008:rosa-2008 01/07/09 15:48 side 42 frequency domain electromagnetic data, reflection seismic profiles and borehole logs. various kinds of 1-d models and 2-d models resulting from inversion of electrical and electromagnetic data are also saved, securing an immediate use of the results. all information about data acquisition, data processing and inversion can be stored, which facilitates reprocessing of data and makes the inversion and interpretation of data transparent. geus also hosts another database (jupiter; http://jupiter. geus.dk) for borehole data. jupiter contains information on, for example, geological and lithological descriptions, groundwater level and water quality observations. both the jupiter and gerda databases have web-based graphical user interfaces, where any user can search for and download data free of charge. geophysical data are handled from data processing to ge ological interpretation in an integrated system formed by the gerda and the jupiter databases and two software packages, the aarhus workbench and the geoscene3d in combination with a geological model database hosted at geus (fig. 2; møller et al. in press). the aarhus workbench (hydro geophysics group 2007b) has modules for handling, processing, inverting, interpreting and visualising electrical and electromagnetic data, all combined on a common gis platform and a common database. the aarhus workbench enables anybody to work with the geophysical data in the gerda database without having to know the complicated data model of gerda or to be able to carry out a database query. by using the gis platform at the aarhus workbench it is easy to produce various types of maps compiled from the geophysical data. the different maps are entered into the 3-d visualisation and modelling tool geoscene3d (i-gis, http://www.i-gis.dk) together with all the geophysical data stored in gerda and the borehole information stored in jupiter, and the geophysical data are ready to be used in the geological modelling process carried out in geoscene3d. an example of the strength of the integrated data handling system is illustrated for a 50 × 60 km2 area in eastern jylland (fig. 3). large parts of this area are covered by tem soundings (c. 83 000 soundings), collected during more than 90 mapping campaigns (fig. 3b) and with five different tem methods (fig. 3c) over a time span of more than ten years. figure 3a shows a map of the surface of the deepest low-resistive model layer based on interpretation of all the tem soundings in the area. the deepest low-resistive model layer represents palaeogene clay deposits except in the north-eastern corner, where it represents salty pore water in danian limestone. the most prominent features found in the area are a large number of buried valleys incised into the palaeogene clay deposits. the buried valleys show no direct correlation to the overall topography. even though the data have been acquired by different companies, with different instruments and methods, and at different times, the data can be combined without showing any discrepancies at survey borders. concluding remarks geophysical measurements play an important role in the national groundwater mapping programme and have contributed significantly to the mapping of aquifers in den mark. in heterogeneous regions the data density needs to be high in order to provide acceptable mapping results. geo physical methods like tem/skytem and electrical methods can provide sufficient data density and reflection seismic profiles can resolve internal structures in specific areas in combination with detailed borehole information such as litho logical descriptions, geophysical logs, data on water chem43 gerda jupiteraarhus workbench data processing raw electrical and electromagnetic data inversion preparation of data to gerda 3-d visualisation data quality control geological modelling geological models hydrostratigraphical models groundwater models borehole information location lithology waterlevel water chemistry wenner seismic data, borehole logs cves schlumberger ip tem paces skytem em3x hem seismics borehole logs 1-d models 2-d models reprocessing re-interpretation visualisation on maps, profiles advanced data analysis advanced data interpretation geoscene 3d modeldb hydrogeological data fig. 2. sketch of the integrated system of databases and program packages handling geophysical data and geological modelling. the arrows show the flow of data between the geophysical database gerda, the borehole database jupiter, the aarhus workbench program package, the geoscene3d visualisation and modelling tool and the geological model database modeldb. rosa_2008:rosa-2008 01/07/09 15:48 side 43 44 istry and hydraulic parameters. these data form the basis for detailed hydrogeological models. an integrated data hand ling system makes it possible to merge geophysical data acquired over long periods by different companies with different instruments. this is of great value for future mapping and administrative purposes. references dahlin, t. 1996: 2d resistivity surveying for environmental and engineering applications. first break 14, 275–283. hydrogeophysics group 2007a: guide to processing and inversion of skytem data. version 1.2. århus: department of earth sciences, uni ver sity of aarhus. http://www.hgg.geo.au.dk/hggsoftware/work-bench/ workbench_skytem.pdf. hydrogeophysics group 2007b: aarhus workbench a-z reference, ver sion 2.2. århus: department of earth sciences, university of aarhus. http://www.hgg.geo.au.dk/hggsoftware/workbench/workbench_az_reference.pdf. jørgensen, f., lykke-andersen, h., sandersen, p.b.e., auken, e. & nør mark, e. 2003: geophysical investigations of buried quaternary valleys in denmark: an integrated application of transient electromagnetic soundings, reflection seismic surveys and exploratory drillings. journal of applied geophysics 53, 215–228. jørgensen, f., kristensen, m., højberg a.l., klint, k.e.s, hansen, c., jordt, b.e., richardt, n & sandersen, p. 2008: opstilling af geologiske modeller til grundvandsmodellering. geo-vejledning 3, 176 pp. copen hagen: geological survey of denmark and greenland. møller, i., søndergaard, v.h., jørgensen, f., auken, e. & christiansen, a.v. in press: integrated management and utilisation of hydrogeophysical data on a national scale. near surface geophysics. rasmussen, e.s., vangkilde-pedersen, t. & scharling, p. 2007: prediction of reservoir sand in miocene deltaic deposits in denmark based on highresolution seismic data. geological survey of denmark and greenland bulletin 13, 17–20. sørensen, k. 1996: pulled array continuous electrical profiling. first break 14, 85–90. sørensen, k.i. & auken e. 2004: skytem – a new high-resolution helicopter transient electromagnetic system. exploration geophysics 35, 191–199. sørensen, k.i., auken, e., christensen, n.b. & pellerin l. 2005: an integrated approach for hydrogeophysical investigations. new technologies and a case history. in: butler, d. (ed.): near-surface geo physics part ii, seg investigations in geophysics series 13, 585–603. tulsa: society of exploration geophysicists. thomsen, r., søndergaard, v.h & sørensen, k.i. 2004: hydrogeologi cal mapping as a basis for establishing site-specific groundwater protection zones in denmark. hydrogeology journal 12, 550–562. vangkilde-pedersen, t., dahl, j. f. & ringgaard, j. 2006: five years of ex perience with landstreamer vibroseis and comparison with conventional seismic data acquisition. proceedings of the 19th annual sageep sym posium on the application of geophysics to engineering and en vironmental problems, seattle, usa, 2–6 april, 2006. (published on cdrom, 1086–1093). authors’ address geological survey of denmark and greenland, lyseng allé 1, dk-8270 højbjerg, denmark. e-mail: ilm@geus.dk fig. 3. data coverage and results from an area in eastern jylland. for location see fig. 1. a: map showing the elevation of the surface of the deepest low-resistive layer in the area relative to sea level. b: the data come from 94 different mapping projects (shown by different colours). c: five diffe rent tem methods were used to produce map a. århus 10 km –200 –100 500 tem 40 hmtem 1 hmtem 2 patem skytem ca b elevation (m) rosa_2008:rosa-2008 01/07/09 15:48 side 44 geological survey of denmark and greenland bulletin 12, 46-57 46 sharp and characterised by prominent shifts on both the sonic and density logs (figs 29, 33). it is often difficult to identify the boundaries on the gamma-ray log alone. depositional environment. although the sandstones of the rind member were deposited from highly concentrated gravity flows, their present appearance largely records postdepositional liquefaction and fluidisation processes. age. thanetian. correlation. the rind member may be contemporaneous with parts of the lithologically dissimilar østerrende clay encountered in the storebælt region (fig. 1), with sandstones in the higher parts of the heimdal formation (hardt et al. 1989) and with the upper balmoral sandstone of the mey sandstone member of knox & holloway(1992).however, therindmember is not contiguous with those sandstone units and has a different source area. sele formation history. the sele formation was established by deegan & scull (1977) for the dark grey to greenish grey, laminated and carbonaceous, tuffaceous, montmorillonite-rich shales and siltstones that overlie the non-laminated and nontuffaceous shales of the lista formation in some areas, or arenaceous sediments belonging to a variety of different units in other areas. the original definition of the sele boundary is followed herein. this implies that the base of the sele formation is located at the base of the “laminated tuffaceous shales” that overlie the “non-laminated, nontuffaceous shales” of the lista formation (deegan & scull 1977; see boundaries section under the lista formation for further details). sandstones occur in the sele formation in the danish sector; these are established as a new member, the kolga member. typewell.british sectorwell21/10-1,2131–2100m mdkb. bue mb rind mb vile mb ve mb balder fm horda fm lark fm sele fm lista fm våle fm 1800 1900 2000 2004.8 2066.3 2100 m sandra-1 gr sonic neutron/density fig. 33. sandra-1, reference well for the rind member. black bar shows cored section. 47 danish reference wells. siri-1, 2072.6–2047.5 m mdkb (fig. 14; plates 1, 4); tabita-1, 2958.8–2941.4 m mdkb (fig. 34; plate 2). distribution and thickness. the sele formation is recognised from a large number of north sea wells and it has a basinwide distribution. in the danish sector, the thickness varies from 5 to 54 m (fig. 35). lithology. the sele formation consists of medium to dark grey, brownish or black laminated mudstones. thin tuff layers occur in the upper part of the formation. it contains three or more well-laminated intervals where dark mudstone beds alternate with lighter coloured mudstone beds. the well-laminated intervals are enriched in organic material resulting in a high gamma-ray response, primarily due to increased uranium content. the most organic-rich, and often darkest, most well-laminated interval is found in the basal part of the formation (fig. 36). the mudstones of the sele formation show an overall upward increase in the silt fraction. in the upper half of the formation, the mudstones may be interbedded with thin, very fine-grained sandstone laminae and thin sandstone beds (fig. 37). the sandstone beds are up to 12 cm thick, normally graded and display parallel lamination. locally, and dominantly in the upper part of the formation, graded tuff laminae less than 1 cm thick are present. in cores, the tuff laminae have a light purple colour. small calcite concretions are present, but rare. in the siri canyon, the sele formation is interbedded with sandstones or it grades upwards into a succession of thinly interbedded sandstones and mudstones. thin sandstone intrusions occur, but only in the lower part of the formation. log characteristics. the sele formation is characterised by high gamma-ray readings throughout, with a number of gamma-ray peaks. on the gamma-ray log, the base of the sele formation is generally marked by a conspicuous upward shift to consistently higher gamma-ray readings than those of the underlying bue member (figs 38, 39). in most wells, a pronounced gamma-ray peak follows a short distance above the base of the sele formation (e.g. augusta-1 and e-8; fig. 38). in wells to the north and west of the danish sector, the stratigraphic distance between the shift to higher gamma-ray readings at the base of the sele formation and the gamma-ray peak is considerably greater (e.g. in the norwegian well 2/7-1; fig. 25). in some wells in the danish sector (and in most siri canyon wells), the basal high gamma-ray interval is missing and the base of the sele formation is marked by the pronounced gamma-ray peak (e.g. cleo-1 and nini-3; figs 38, 39). boundaries. the lower boundary is characterised by a change from the light to dark grey and greyish black mudstones with thin sandstone laminae of the bue member (lista formation), to dark grey to black well-laminated mudstones without sandstone laminae of the sele formation (fig. 36). the upper boundary is at the base of the balder formation. subdivision. knox & holloway (1992) suggested an informal threefold subdivision of the sele formation based fig. 35. isochore map of the sele formation in the study area. the positions of the two danish reference wells, siri-1 and tabita-1, are indicated in the figure. siri-1 tabita-1 10 20 30 40 50 thickness (m) sele formation 25 km fig. 34. tabita-1, danish reference well for the sele formation. horda fm balder fm sele fm lista fm vile mb ve mb bue mb våle fm chalk gp 2900 3000 m tabita-1 2941.4 2958.8 gr sonic 48 in figs 24, 26 and 38), whereas other lithological changes are more subtle. the subdivision is outlined below. unit s1 this unit comprises a lower subdivision s1a and an overlying subdivision s1b. subdivision s1a is identical to the bue member of the lista formation and the stratigraphic interval from the base of the sele formation (sensu deegan & scull 1977) up to the base of the lowermost conspicuous gamma-ray peak within that formation (fig. 38). that peak is associated with relatively low sonic values. subdivision s1b has its base at the gamma-ray peak and its top at the base of the next gamma-ray peak. the gamma-ray response decreases up through s1b. unit s1 consists of brownish grey to dark grey and black, well-laminated mudstones. unit s2 the base of the unit is at the base of the second gamma-ray peak. this peak can be differentiated from the gamma-ray peak at the base of s1b by its association with high sonic values. unit s2 can be divided into a lower subdivision (s2a) characterised by a relatively high gamma-ray response level and an overlying subdivision (s2b) with a lower gamma-ray response (fig. 38). the two subdivisions are separated by a gammaray low. unit s2 consists of light grey to brownish dark grey, laminated to well-laminated mudstones. both the lamination and the colour of the two subdivisions are very similar in the cores encountered in this study and it is almost impossible to distinguish the two subdivisions on lithology alone. knox & holloway (1992) observed tuff layers in the basal part of unit s2b in the british wells; tuffs were not observed in the danish wells.. unit s3 this unit is characterised by high and increasingupwards gamma-ray values. the base of the unit is defined by a sharp increase in gamma-ray readings (fig. 38). there is no significant colour difference between the mudstones of units s2 and s3 in the danish wells, but lamination seems to be better developed in unit s3 than in unit s2. tuff layers were observed in unit s3 in the british wells studied by knox & holloway (1992). similar tuff layers have been observed in wells from the siri canyon, and may further be used to distinguish unit s3 from the upper part of unit s2. 0 10 20 30 40 50 60 70 80 90 100 cm 2888 m 2889 m augusta-1 2890 m sele fm bue m b fig. 36. core photographs of the lista–sele formation boundary interval in the augusta-1 well. the upper part of the bue member (lista formation) consists of mudstones with thin sandand siltstone laminae superficially resembling mudstone-in-mudstone lamination. at 2890.35 m, the bue member is overlain sharply by the laminated mudstones of the sele formation. depths are core depths. on the gamma-ray log signature. this subdivision can also be recognised on petrophysical logs from most wells in the danish sector (the subdivision is shown in five wells 49 0 10 20 30 40 50 60 70 80 90 100 cm 1689 m 1690 m nini-3 1692 m 1693 m the sele formation includes a sandstone unit (kolga member, new) in the danish north sea sector. macroand ichnofossils. fish scales and skeletal fragments are common in cores from the sele formation. bioturbation is very rare, but chondrites ispp. has been observed locally. microfossils and palynomorphs. benthic foraminifers are rare in the sele formation. the lo of an acme of the dinoflagellate genus apectodinium and the coeval lo of the shortranged a. augustum mark a level at, or a few centimetres above the base of the sele formation. the ho of a. augustum is located in the lower part of the sele formation. the ho of an influx of the dinoflagellate cerodinium wardenense marks a level in the upper part of the sele formation. the ho of an influx of the diatoms fenestrella antiqua and coscinodiscus morsianus is located in the uppermost part of the sele formation. throughout, the formation contains abundant spores and pollen, in particular pollen of the genus inaperturopollenites. depositional environment. the mudstones of the sele formation represent a mixture of pelagic fallout and dilute, low-density mud turbidites. the well-laminated character of the sediment, the high content of organic material and uranium, and the general lack of trace fossils and benthic foraminifers indicate starved sedimentation under dysoxic to anoxic bottom conditions. common diatoms indicate a high nutrient level in the water mass. the tuffs of the sele formation are evidence of extensive volcanism in the region. the significant depauperation of the benthic microfaunas during the deposition of the sele formation was most likely caused by isolation of the north sea basin (schmitz et al. 1996). the restriction and isolation of the basin was the result of a sea-level fall, possibly combined with (or caused by) tectonic uplift to the north-west (knox et al. 1981). based on microfossils, the palaeoenvironment has been suggested to represent an upper bathyal setting with a palaeodepth estimate of around 300 m (mitlehner 1996). the palynomorph assemblage indicates a marine environment characterised by a massive influx of terrestrial palynomorphs. age. sparnacian (sensu aubry et al. 2003) – early ypresian, with the lowermost level possibly of thanetian age. correlation. the sele formation corresponds to the haslund member of the ølst formation (heilmann-clausen et al. 1985). its upper part correlates with the haslund member-equivalent diatomitic knudeklint member of the fig. 37. core photographs from the nini-3 well showing the upper, arenaceous part of the sele formation. the interval belongs to the s2a subunit of the sele formation (see text for further explanation). stratigraphic position of the figured interval is shown in fig. 39. depths are core depths. fur formation in north-west jylland (danielsen & thomsen 1997). the lower boundary of the sele formation correlates with the boundary between the østerrende clay and the haslund member onshore denmark. 50 2760 2740 m 2780 s3 s2b s2a s1b s1a 2900 cleo-1 augusta-1 e-8horda fm balder fm sele fm lista fm bue mb ve mb 2880 2040 2860 2020 m m gr sonic gr sonic gr sonic cleo-1 e-8 the lower part of the sele formation, consistingof laminated, dark grey to black mudstones, correlates with the lithologically very similar, 15 m thick informal unit stolle klint clay that constitutes the lower part of the haslund member, onshore denmark (heilmann-clausen 1995). this unit is known from throughout the north sea (hardt et al. 1989; knox & holloway 1992) and constitutes most or all of the sele s1b unit of knox & holloway (1992). kolga member new member history. the kolga member consists of sandstone deposits within the sele formation. these sandstones were previously recognised by a stratigraphic working group at statoil norway in the mid-1990s and were informally referred to the hermod formation of hardt et al. (1989). derivation of name. after the goddess kolga. type well. danish sector well siri-3, 2066.4–2036.1 m mdrt (fig. 24; plate 4). reference well. danish sector well nini-3, 1717.2–1700.4 m mdrt (figs 26, 39; plate 4). distribution and thickness. the kolga member has a restricted distribution in the siri canyon in the northern part of the danish sector (fig. 20c). it reaches a thickness of up to 30 m. lithology. the member consists primarily of fine-grained to very fine-grained, olive-green to greenish grey, well-sorted, quartz-rich sandstones (fig. 39). rounded and translucent quartz grains dominate the mineralogical assemblage, but the content of glaucony grains is high (15– 25%). mica and small pyrite concretions are present in small amounts. locally, the sandstones are partly cemented by calcite and chlorite. the member usually includes one thick unit composed of amalgamated sandstone beds and a number of thinner sandstone beds interbedded with mudstones that are lithologically comparable to the sele formation mudstones described above. log characteristics. the kolga member is clearly defined on the gamma-ray log by a blocky pattern with intermediate values. this pattern differs from the high gamma-ray readings that normally characterise the lower sele formation. the kolga member may show a gradual upward decrease in gamma-ray response in its lower part (e.g. in the well nini-3; figs 26, 39). however, this does not reflect grain-size change, judging from core examination. the density log shows a blocky pattern with low density fig. 38. correlation diagram of the sele formation showing the sele units s1–3 of knox & holloway (1992). 51 v v v sele fm lista fm kolga mb bue mb grsonic 1700 1690 1710 1720 1700 1710 1720 nini-3 clay si. vf. f. m. sand c. vc. log depth core depthfig. 39. core log of the sandstonedominated kolga member encased in the sele formation mudstones in the nini-3 well. for legend, see fig. 9. intervals marked by grey bars in the core depth column are shown as core photos in fig. 37. values for the kolga member sandstones and relative high values for the interbedded mudstones (fig. 24). boundaries. the boundaries between the sandstones of the kolga member and the mudstones of the sele formation are sharp and characterised by prominent shifts on the gamma-ray and sonic logs (figs 24, 26). microfossils and palynomorphs. the kolga member is characterised by an abundance of apectodinium spp., including a. augustum. depositional environment. the sandstones of the kolga member were deposited from highly concentrated gravity flows, although the present character of the kolga member mainly reflects postdepositional liquefaction and fluidisation processes. primary sedimentary structures are common in the kolga member in some wells (e.g. sandra-1),however, indicatingthat thememberhasexperienced less postdepositional remobilisation in certain areas. age. sparnacian (sensu aubry et al. 2003) possibly including the latest thanetian. correlation. the kolga member is contemporaneous with parts of the sele formation onshore denmark. it possibly correlates with the forties sandstone member (knox & holloway 1992) in the central graben and with the teal and skadan sandstone members (knox & holloway 1992) in the southern viking graben. however, the kolga member is not contiguous with those units and has another source area. fur formation history. the fur formation is a marine diatomite with numerous ash layers. it was formally established by pedersen & surlyk (1983) with a type section in the coastal cliff knudeklint on the island of fur, denmark. its lower boundary was revised by heilmann-clausen et al. (1985). the characteristic lithology of the fur formation was subsequently recognised by thomsen & danielsen (1995; danielsen & thomsen 1997) in cuttings samples from three offshore wells located in the north-eastern part of the danish sector of the north sea, as well as in one well in the norwegian sector. 52 350 450 m k-1 gr sonic balder fm horda fm fur fm sele fm lista fm våle fm chalk group 398.9 402.2 400 m 485.5 697.1 795.3 810.7 500 600 700 800 900 dufa mb horda fm balder fm lark fm sele fm fur fm lista fm våle fm chalk gp inez-1 gr sonic type section. the coastal cliff knudeklint, the island of fur, onshore denmark (for location map, see pedersen & surlyk 1983). reference sections. silstrup south cliff, skarrehage, feggeklit, harhøj, stolleklint (for location maps, see pedersen & surlyk 1983). danish reference wells. danish sector wells k-1, 402.2– 398.9 m mdkb (fig. 40; plate 5); inez-1, 810.7–795.3 m mdkb (fig. 41; plates 1, 5). distribution and thickness. onshore denmark, the fur formation is distributed in a limited area in north-west jutland. it is c. 61 m thick in its type section. offshore, the formation occurs in a belt stretching from the northwestern coast of jutland, continuing into the norwegian sector parallel to the southern coast of norway (thomsen & danielsen 1995 text-fig. 6; fig. 42). it reaches a thickness of 15.4 m in the inez-1 well, 7.9 m in the c-1 well and 3.3 m in the k-1 well (figs 40–42). lithology. the lithology of the fur formation was described from its onshore exposures by pedersen (1981) and pedersen & surlyk (1983). it is a clayey, porous, dark grey diatomite with numerous volcanic ash layers. diatom frustules constitute 65 wt% of the rock, clay particles constitute 35 wt% (pedersen 1981). tests of coscinodiscus spp. and stephanopyxis are the major constituent of the diatomite fraction (thomsen & danielsen 1995). fine lamination is the primary sedimentary structure, but at some fig. 41. inez-1, reference well for the fur formation and type well for the dufa member. fig. 40. k-1, reference well for the fur formation. 53 25 km k-1x 3.3 15.4 7.9 inez-1 c-1 57°00' 6°00' 8°00' 56°00' 55°00' ? ? ? ? fur formation levels the lamination has been destroyed by bioturbation. the ash layers are black, graded and consist of volcanic glass particles. individual layers range from 1–20 cm in thickness but are fairly uniform in thickness over a limited area (pedersen & surlyk 1983). the diatomite recognised in cuttings samples from north sea wells by thomsen & danielsen (1995) is a similar lithology to that of the onshore sections studied by pedersen (1981) and pedersen & surlyk (1983). log characteristics. the fur formation is identified by the combination of a low gamma-ray response and low sonic readingswithinan interval of higher sonic readings characterising mudstones below and above (thomsen & danielsen 1995; danielsen& thomsen1997;figs40,41;plates 1, 5). boundaries. in the c-1 and k-1 wells, the fur formation is enveloped by the balder formation. in inez-1, the formation is bounded by the sele and balder formations (figs 40, 41; plates 1, 5). it should be noted, however, that thomsen & danielsen (1995 text-fig. 4) and danielsen & thomsen (1997 fig. 5) placed the fur formation entirely within the sele formation in the k-1 well. the discrepancy between the interpretation herein and that of the former authors is due to different interpretations of the position of the balder–sele boundary in the well. the boundary of the fur formation with the balder and sele formations is characterised by a change from laminated or structureless diatomite with ash layers to the dark mudstones of the balder and sele formations. this lithological change is reflectedonthe sonic logbyan abrupt increase in velocity (figs 40, 41; plates 1, 5). subdivision. onshore denmark, the fur formation is divided into the lower, laminated knudeklint member that contains relatively few, widely spaced ash layers and the upper, mainly structureless silstrup member with numerous ash layers (pedersen & surlyk 1983). macroand ichnofossils.themacrofossil assemblage described from onshore exposures of the fur formation encompasses fish, birds, turtles, snakes, starfish, shellfish, snails, mussels, crabs, pteropods, insects, fossil wood, leaves and fructifications (bonde 1966, 1979, 1987, 2003; pedersen 1981; pedersen & surlyk 1983; kristoffersen 2001). ichnofossils from onshore exposures include planolites ispp., teichichnus ispp., chondrites ispp. and taenidium ispp. (pedersen & surlyk 1983). microfossils and palynomorphs. diatom frustules are rockforming in the fur formation (pedersen 1981). silicofig. 42. distribution map of the fur formation with formation thickness (m) indicated for three wells. flagellates are present (perch-nielsen 1976) whereas calcareous microfossils are absent. the formation contains abundant dinoflagellates and sporomorphs (hansen 1979; heilmann-clausen 1982; willumsen 2004). depositional environment. the deposition of the fur formation diatomites took place in a long, narrow zone under upwelling conditions (bonde 1974, 1979). the upwellingwas controlled by northerly winds (bonde 1974, 1979) or it may have been created by a combination of bottom currents and bottomtopography (pedersen& surlyk 1983). age. early ypresian. correlation. onshore denmark, the knudeklint and silstrup members are largely contemporaneous with the upper part of the haslund member and the overlying værum member of the ølst formation (heilmannclausen et al. 1985), respectively, and correlate with the 54 upper part of the sele formation and the lower unit b1 (see below) of the balder formation (heilmann-clausen 1995; knox 1997 fig. 3). balder formation history. deegan & scull (1977) established the balder formation for the succession of variegated, fissile and laminated shales with interbedded tuff layers that lie between the sele and horda formations. type well. norwegian sector well 25/11-1, 1780–1705 m mdkb. danish reference wells. mona-1, 2945.0–2930.8 m mdkb (fig. 43; plate 1). siri-3, 2016.8–1998.8 m mdrt (fig. 24; plate 4). distribution and thickness. the balder formation extends over most of the central and northern north sea. in the danish sector, it reaches a thickness of more than 20 m in the siri-3 and frida-1 wells on the western part of the ringkøbing–fyn high (fig. 1) and 20 m in gwen-2 in the northern part of the danish sector of the central graben. the balder formation thins to less than 5 m towards the south-west and to less than 10 m in the eastern part of the danish sector of the north sea. the balder formation is lacking in the danish well s-1 (michelsen et al. 1998). an isochore map of the balder formation is shown in fig. 44. lithology.thebalderformationis composedof laminated, dominantly grey, fissile shales with interbedded dark and light grey, purple, buff and green sandy tuffs (fig. 45). the tuffs are normally graded and less than 5 cm thick. locally the tuff beds are slumped. the tuff layers may be cut by irregular, vertical, calcite-filled cracks up to 20 cm long (fig. 45). similar cracks have been reported from the balder formation in the grane field, norwegian sector of the north sea (haaland et al. 2000). sandstone beds, interpreted as intrusive sandstone bodies, occur locally in the balder formation. log characteristics. the balder formation is characterised by a relatively high gamma-ray values in its lower and higher parts, but shows low values in its middle part. the change in gamma-ray response is normally gradual, but relatively steep. the gamma-ray motif is mirrored by a gradual increase in sonic readings commencing at the formation base, culminating at or slightly below the level of minimum gamma-ray values in the middle part of the fig. 44. isochore map of the balder formation in the study area. the positions of the two danish reference wells, mona-1 and siri-3, are indicated in the figure. siri-3 mona-1 balder formation 5 10 15 20 thickness (m) 25 km fig. 43. mona-1, danish reference well for the balder formation. 3000 2900 m 2930.8 2945.0 horda fm balder fm sele fm lista fm våle fm vile mb ve mb bue mb chalk gp mona-1 gr sonic 55 2013 m siri-3 2012 m 2014 m 2015 m 0 10 20 30 40 50 60 70 80 90 100 cmformation, followed by a gradual decrease towards the top of the formation where the lowest sonic reading is reached. the gamma and sonic motifs together create a characteristic barrel-shaped log pattern (e.g. figs 24, 33, 38, 43; plates 1–5). boundaries. in general, the boundary with the underlying sele formation is gradational, although it can be sharp in some wells. where gradational, it is placed where the tuff layers become prominent (e.g. fig. 45). on petrophysical logs, the lower boundary is identified at a significant upward decrease in gamma-ray response accompanied by an increase in sonic readings (e.g. figs 24, 33, 38, 43). the upper boundary is at the base of the horda formation. subdivision. knox & holloway (1992) subdivided the balder formation into a lower, laminated and tuff-rich unit (b1) and a poorly laminated upper unit (b2). this subdivision can be recognised in a number of danish wells west of cecilie-1, but the b2 unit seems to be absent from danish north sea wells north-east of, and including, the cecilie-1 well. macroand ichnofossils. macrofossils have not been observed. the balder formation is non-bioturbated to moderately bioturbated. ichnofossils comprise chondrites ispp., phycosiphon ispp., planolites ispp. and thalassinoides ispp. microfossils and palynomorphs. the sele–balder boundary interval is characterised by the ho of common fenestrella antiqua and coscinodiscus morsianus (both diatoms). this event is located in the uppermost part of the sele formation but may be used as a biostratigraphic guide to locate the boundary. the diatom fenestrella antiqua characterises the balder formation and has its ho at the formation top. the dinoflagellate deflandrea oebisfeldensis shows an acme at the top of the balder formation. as observed in the underlying sele formation, the balder formation contains high numbers of spores and pollen, in particular pollen of the genus inaperturopollenites spp., and the top of the balder formation is marked by the ho of common representatives of that genus. in contrast to the overlying horda formation, calcareous benthic foraminifers are virtually absent in the balder formation. depositional environment. a restricted marine palaeoenvironment at upper bathyal depths with dysoxic to anoxic bottom conditions is suggested for the balder formation. this is based on the scarcity of calcareous microfossils and agglutinated foraminifers combined with common to abundant siliceous microfossils, especially diatoms. the fig. 45. core photographs of tuffaceous balder formation mudstones from the siri-3 well. the tuff layers are seen as light coloured, graded intervals (e.g. at 2012.42–2012.40 m). the boundary with the underlying sele formation is placed where tuffs become common, at 2015.5 m (large arrow); two tuff layers may be seen in the uppermost sele formation, at 2015.66 and 2015.90 m (small arrows). two small, lightning-shaped cracks are seen at 2015.4 m. depths are core depths. 56 2363.5 2930.8 1598.3 horda fm h2 h1 h3 l2 l3 l4 lark fm nordland gp balder fm sele fm lista fm våle fm chalk gp 1600 1700 1800 1900 2000 2100 2200 2300 2400 2500 2600 2700 2800 2900 3000 m mona-1 gr sonic fig. 46. mona-1, danish reference well for the horda and lark formations. the figure shows the tripartite subdivision of the horda formation and the l2–4 units of the lark formation. the l1 unit is absent in the mona-1 area. presence of abundant terrestrial palynomorphs further supportsa restricted,marginalmarinedepositional setting. a petrographic and geochemical study of the balder formation in the grane field, norwegian north sea sector, shows that the tuffs can be classified as representing sub-alkaline basalts and basaltic andesites of intra-plate origin (haaland et al. 2000). the tuffs are similar to the contemporaneous lower basalts in east greenland, the rockall trough and the middle series of the faeroe islands, all linked to the opening of the north atlantic (haaland et al. 2000). the volcanic phase took place at 55–52 ma. age. early ypresian. correlation. although unit b2 of knox & holloway (1992) is apparently lacking in wells in the north-eastern part of the danish sector, both units b1 and b2 can be correlated with strata onshore denmark, although unit b2 is very thin. unit b1 corresponds to the lithologically similar værum member of the widespread ølst formation onshore denmark and with the diatomaceous silstrup member of the fur formation in north-west jylland (knox 1997 fig. 3). the lower boundary of the værum and silstrup members is placed at ash layer no. +1 in the tephrachronology of bøggild (1918). the ash chronology has not been identified in the type section of the balder formation and precise correlation with the lower boundary of the værum and silstrup members is therefore uncertain. however, judging from the abundance of thick ash layers in the balder formation and the scarcity of ash layers in the underlying sele formation, it is likely that the base of the balder formation approximately correlates with ash layer no. +1, i.e. with the base of the værum and silstrup members. according to knox (1997 fig. 3), unit b2 probably correlates with the knudshoved member of the røsnæs clay formation (heilmann-clausen et al. 1985). this member has a very restricted distribution in north-west jylland where it overlies the silstrup member of the fur formation. the knudshoved member consists of a lower dark grey, pyritic clay unit rich in pyritised diatoms, and an upper greenish clay unit (heilmann-clausen et al. 1985). only a few, thin volcanic ash layers are present in the member (håkansson & sjørring 1982). based on lithological comparison, it is suggested that at least the lower, 57 pyritic part of the knudshoved member may correlate with theupper, tuff-poorunitb2 of the balder formation. stronsay group knox & holloway (1992) replaced the hordaland group of deegan & scull (1977) with two new groups: the stronsay group succeeded by the westray group (fig. 3). the two groups together comprise the light grey, green and brown coloured, soft, fissile, marine shales with thin limestone streaks that overlie the rogaland group and underlie the nordland group. these groups each contain two formations, one representing sandy shelf lithofacies and the other representing basinal mudstone lithofacies. in the central north sea, and in the danish sector, the stronsay group is represented by its mudstone facies, the horda formation (knox & holloway 1992). sandstone units of varying thickness occur at many levels in the stronsay and westray groups along the basin margin in the norwegian and british sectors, and many of these have been defined as formations or members (deegan & scull 1977; hardt et al. 1989; knox & holloway 1992). a sandstone unit also occurs in the horda formation on the ringkøbing–fyn high in the danish sector and is described here as a new member (hefring member). horda formation history. knox & holloway (1992) established the horda formation for the greenish grey basinal mudstone facies of their stronsay group that overlies the grey tuffaceous mudstones of the balder formation and underlies the greenish grey to brown mudstones of the lark formation (knox & holloway 1992). type well. british sector well 22/1-1a, 2379.5–1992 m mdkb. danish reference wells. mona-1, 2930.8–2363.5 m mdkb (fig. 46); siri-1, 2037.9–1916.5 m mdkb (fig. 47). distribution and thickness. the horda formation extends over the central and northern north sea and is present in 1916.5 2037.9 horda fm chalk gp lark fm balder fm sele fm lista fm våle fm 1900 2000 2100 2200 m siri-1 gr sonic rind mb idun mb tyr mb vile mb ve mb bue mb bue mb horda formation siri-1 mona-1 floki-1 100 200 300 400 500 600 700 800 900 thickness (m) 25 km fig. 47. siri-1, danish reference well for the horda formation. black bars show cored sections. fig. 48. isochore map of the horda formation in the study area. the positions of the two danish reference wells for the horda formation, mona-1 and siri-1, are indicated on the map. the position of floki-1, the type well for the hefring member, is also indicated. geological survey of denmark and greenland bulletin 17, 2009, 77-80 the third galathea expedition (galathea 3) left copenhagen in august 2006 for a circumnavigation of the globe with the aim of conducting more than 70 scientific programmes en route. the first geological programme took place in south greenland and included sampling of sediment cores and seismic profiling. the aim of the study is to obtain detailed knowledge about holocene climate changes and the glacio marine history. a number of cores with high sedimentation rates were collected near narsaq in south greenland (fig. 1). analyses of the cores elucidate the midto late holocene climatic and environmental evolution of the area, which is highly influenced by the dynamic nature of the greenland ice sheet and changes of the north atlantic climate. two major ocean currents, the cold east greenland current and the warmer (atlantic water) irminger current influence the deep fjords of the region, and the two currents are in turn influenced by both polar and lower latitude climate changes. evidence from ice cores and marine and lacustrine records in the north atlantic region show a general climatic cooling from the midto late holocene as a response to decreasing summer insolation. about 5000 years ago, a transition occurred from the holocene thermal maximum to the neoglacial (dahl-jensen et al. 1998). the cooling trend shows different timing and amplitudes in different parts of the north atlantic region (kaufman et al. 2004), and an apparent counterphase between the climatic conditions of western greenland and north-western europe during certain time periods has recently been suggested (seidenkrantz et al. 2008). the present study investigates climate changes in south greenland in relation to the previously established pattern of holocene palaeoceanographic and atmospheric changes as known from other north atlantic palaeoclimatic studies. study area and methods glacially eroded and over-deepened fjords reaching depths of 600–700 m dissect south greenland around qaqortoq and narsaq (fig. 1). sediment cores were collected in bredefjord and narsaq sund (fig. 1) using a gravity corer and a box corer at water depths of 270–670 m. a gravity corer with a 750 kg lead weight was used to collect sediment cores up to 6 m long with a diameter of 12 cm. surface sediments were sampled with a cylindrical 30 cm diameter box corer. all cores were subsampled for analyses of microfossils and age determination using accelerator mass spectrometry 14c dating. the cores were split lengthwise, and magnetic susceptibility was 77 holocene climate variability in southern greenland: results from the galathea 3 expedition niels nørgaard-pedersen, naja mikkelsen, majken djurhuus poulsen and aaju s. simonsen © geus, 2009. geological survey of denmark and greenland bulletin 17, 77–80. available at: www.geus.dk/publications/bull fig. 1. map of south greenland showing the main current systems and locations of investigated cores. the waters of the east greenland current and the warmer and more saline irminger current (deeper than 200 m) form a stratified water column along the coast of southern greenland. bredefjord qassimiut lobe 800–2500 m 200–800 m narsaq sund 1250 2000 narsaq 50 km qaqortoq narsarsuaq irminger current east greenland current ic eg c ga3-2 ga3-7 ga3-11 ga3-10 iceland greenland 60°n 60°n 44°w 46°w above 200 m rosa_2008:rosa-2008 01/07/09 15:49 side 77 measured in high resolution using a bartington ms2e1 probe. nine samples of benthic foraminifera were used for 14c dating at the leibniz laboratory for radiometric dating and isotope research in kiel, germany, and the box cores were analysed for 210pb and 137cs content at the gamma dating center, department of geography and geology, uni versity of copenhagen. a number of basic sediment parameters were determined including sediment wet-bulk density, water content, dry-bulk density and grain-size distribution. magnetic susceptibility measured at high resolution in general parallels the records of coarse-fraction content, and can therefore be used for correlation and as a high-resolution measure of coarse-fraction influx. in order to estimate the variation in bottom-current velocity, detailed grain-size measurements (range 0.01–1000 µm) were carried out on bulk-sediment samples using a malvern mastersizer 2000 laser particle-size analyser at the department of geography and geology, university of copenhagen. the weighted grain-size means of the sortable silt fraction (10–63 µm) were calculated and used as a proxy for bottom current. the number of lithogenic grains >1 mm (coarse ice-rafted debris, ird) was counted using a dissecting microscope. the rationale for using grains >1 mm is a compromise between having an adequate number of grains and using grains so large that aeolian transport can be considered negligible. the ird content is thus used as a proxy for rock fragments from melting of icebergs or sea ice deposited at the site. the sediment-size fraction 125–1000 µm was separated by dry sieving, and from this fraction 300–400 calcareous benthic foraminifera were picked for analyses. in sediment cores with a fairly uniform sedimentation rate and limited calciumcarbonate dissolution, the abundance of calcareous benthic foraminifera mainly reflects productivity related to food supply. fjord environment and sedimentation records the records from the three gravity cores from bredefjord are dominated by episodic turbidite sedimentation with very high sedimentation rates of up to 1 cm/year (based on a calibrated 14c age from ga3-7 (435 cm depth) of c. 465 years bp) and is under influence from meltwater supply and gla cier calving from the adjacent margin of the greenland ice sheet (fig. 1). for this study, we selected core ga3-2 from the adjacent narsaq sund (fig. 1), which represents a continuous and more slowly accumulating (c. 70 cm/1000 years) sedimentation record covering the last c. 8000 years. the chronology is based on eight 14c ages (fig. 2). the record is characterised by a grain-size spectrum of dominantly suspension-settled mud modulated by variable bottom-current in fluence and ird rain (fig. 3). the sedimentation in this turbidite-sheltered part of the fjord system has been controlled by a number of related processes including influx of sediment-laden meltwater plumes with fine-grained sediment, iceberg calving and coarse-sediment ird transport and strength and changes of fjord circulation/water-mass stratification by irminger current and east greenland current water masses 78 mwp rwpdalia holocene thermal maximumneoglaciation 0 2 4 6 8 10 12 0 1000 2000 3000 4000 5000 6000 7000 8000 0 100 200 300 400 500 600 c ur re nt s tr en gt h 0 100 200 300 400 500 600 20 22 24 26 28 30 0 100 200 300 400 500 600 so rt ab le s ilt m ea n (µ m ) be nt hi c fo ra m s/ g m ag n. s us c. ( 10 -6 s i) ir d > 1m m ( gr ai ns /g ) d ep th ( cm ) age (years bp) age model: calibrated 14c dates (2 sigma age range) fig. 2. the narsaq sund record (core ga3-2). the chronology is based on 14c age determinations of eight benthic foraminifera samples. sedi ment parameters shown: magnetic sus ceptibility with five-point running mean, icerafted debris content (ird >1 mm), weighted mean grain size of sortable silt (bottom current proxy), and abundance of calcareous benthic foraminifera/g dry sediment. the grey columns indicate periods of increased bottom currents during the neo gla ci ation. european historical climate periods (cf. lamb 1995) are indicated: lia, little ice age; mwp, medieval warm period; da, ‘dark ages’; rwp, roman warm period. rosa_2008:rosa-2008 01/07/09 15:49 side 78 penetrating into the fjord system (ribergaard 2007). modified irminger water is found below c. 200 m water depths and east greenland waters above. during the summer a local brackish meltwater-rich layer characterises the uppermost 10–15 m of the water column. the narsaq sund record the lower part of the narsaq sund record, deposited from about 8.0 to 4.8 ka (ka = thousand years before present), represents the holocene thermal maximum and shows a progressive reduction in the supply of coarse-grained material, punctuated by increased influx of ird at about 7.5–6.8 and 6.5–5.7 ka (fig. 2). the interval 5.7–4.8 ka appears to have been characterised by markedly lower amounts of ird, and it is suggested that this could reflect a situation in which several of the present tidewater glaciers terminated on land. the neoglaciation appears to have started at about 4.8 ka, with peak values of ird at about 4.6, 3.6, 2.2, 1.0, 0.7, and 0.5 ka and onward (fig. 2). several of these ird events appear to be associated with enhanced melting and glacier instability during warming episodes. on the other hand, a marked increase in ird is noted at the transition to the little ice age at about 0.55 ka. this appears to correlate with the maximum ice ad vance of the qassimiut lobe (weidick et al. 2004). a threshold may have been reached during this period, with a marked advance of tidewater glaciers to a new equilibrium state leading to increased ird deposition in the fjords. assemblages of calcareous benthic foraminifera show that the bottom of narsaq sund has been dominated by modified irminger water during most of the holocene. bottom-current flow seems to have been quite sluggish before 3.2 ka, whereas the late holocene was characterised by pronounced episodes of increased bottom-current velocity and increased productivity of calcareous benthic foraminifera (figs 2, 3). the first episode at 3.2–2.8 ka shows no relation to ird/ meltwater proxies and the timing at 3.2 ka indicates that the episode may be related to a concurrent major reorganisation of the north atlantic current systems (cf. koç et al. 1993). during two following periods, at c. 2.3–1.7 ka and c. 1.3–0.8 ka, marked increases in bottom-water flow also occurred, and it is suggested that the fjord circulation during these time intervals was enhanced by an increased outflow of meltwater. the timing of the latter two events may imply a link to climate changes during the european roman warm period and the medieval warm period (cf. lamb 1995). radiocarbon dating and palaeoecological studies show that the fjord system became ice-free at 13 to 9 ka and the east greenland current and the irminger current entered the deep and glacially eroded fjords (weidick et al. 2004). during the recession of the greenland ice-sheet margin from the region, large amounts of icebergs characterised the fjords, but with the onset of the holocene thermal maximum at 8–5 ka, the temperature increased and became a few degrees higher than today. during the holocene thermal maximum the greenland ice sheet reached a minimum size. it became increasingly colder in southern greenland after the thermal maximum, and with the onset of colder conditions during the neoglacial period at 4–3 ka the glaciers readvanced. during the little ice age that culminated around ad 1600–1850, the ice margin expanded several kilometres and the greenland ice sheet reached its maximum size after the holocene thermal maximum (weidick et al. 2004). during the late holocene southern greenland subsided 6–8 m, probably partly due to the re-advance of the greenland ice sheet (sparrenbom et al. 2006). at the present time most glaciers in greenland display marked recession. the see-saw puzzle of holocene north atlantic climate recently, seidenkrantz et al. (2008) studied late holocene marine records from west greenland and found a complex relationship between the palaeoclimate of west greenland and that of other parts of the north atlantic region. their study indicated that in west greenland, the european med ieval warm period (c. 1.3–0.8 ka in the greenland record) was characterised by a decreased influence of the warm 79 0.0 0.4 0.8 1.2 1.6 2.0 2.4 2.8 3.2 3.6 4.0 4.4 age (years bp) pa rt ic le s iz e (µ m ) 1000 1000 2000 3000 4000 5000 6000 7000 8000 100 10 1 vol.-% fig. 3. particle-size variation plotted against age for the narsaq sund ga3-2 core. a distinct increase in the modal peak of the grain-size distribution occurred at about 3.2 ka. this suggests an increased bottom-water circulation in narsaq sund, which may be related to changes of the north atlantic current systems or increased meltwater discharge. rosa_2008:rosa-2008 01/07/09 15:49 side 79 irminger current and thus a relatively cool climate, whereas the colder european ‘dark ages’ (c. 1.5–1.3 ka in the greenland record) appear to have been characterised by a higher meltwater runoff and consequently a warmer climate (seidenkrantz et al. 2008). in contrast, marine, lacustrine and ice-core studies from southern greenland (e.g. dahl-jensen et al. 1998; andresen et al. 2004) do not reveal a convincing climate see-saw trend between southern greenland and northern europe. southern greenland is located on the main track of cyclones crossing the north atlantic region on their way towards iceland. therefore the area has potential to provide crucial information on major shifts in the north atlantic atmospheric circulation and general climate regime. the long-term pattern appears to follow climate changes consistent with patterns of east greenland current variability (bond et al. 2001), which are partly related to the dominant north atlantic oscillation (buch 2002). this long-term seesaw pattern of the influence of the irminger current reported by seidenkrantz et al. (2008) is evident in the recent north atlantic record. related changes in air temperature, sea-surface temperature, storm patterns, and sea-ice distribution offer a likely explanation for long-term changes and different regional climate trends in the north atlantic region. there are, however, many unsolved questions in this puzzle, and more high-resolution records are needed from selected sites in order to investigate causes and links between regional climate variations in the north atlantic region. climate change and the demise of the norse norse immigrants settled around ad 985 in south green land where farming communities were established in the eastern settlement during the medieval warm period. the norse society no doubt had to cope with a number of socioeconomic and environmental problems that eventually caused their disappearance after almost 500 years of existence in greenland. one of their problems was the general temperature decrease at the transition from the medieval warm period to the little ice age, discussed by mikkelsen et al. (2008). this transition may have brought their living conditions to a critical point although the manner of the demise of the norse people is still an unsolved question. temperature decrease, increasing storminess and sea-ice increase threatened the existence of the norse people and presumably led them to gradually leave greenland – perhaps for iceland from where their pioneer ancestors originally set out. acknowledgements we thank captain and crew of the danish naval ship vædderen for their help and support during the cruise. we gratefully acknowledge financial support from the bg fund, det kongelige grønlandsfond and the commission for scientific research in greenland. this is publication no. p29 on the scientific outcome of the galathea 3 expedition in 2006–2007. references andresen, c.s., björck, s., bennike, o. & bond, g. 2004: holocene climate changes in southern greenland: evidence from lake sediments. journal of quaternary science 19, 783–795. bond, g., kromer, b., beer, j., muscheler, r., evans, m.n., showers, w., hoffmann, s., lottibond, r., hajdas, i. & bonani, g. 2001: persistent solar influence on north atlantic climate during the holocene. science 294, 2130–2136. buch, e. 2002: present oceanographic conditions in greenland waters. danish meteorological institute, scientific report 02-02, 39 pp. dahl-jensen, d., mosegaard, k., gundestrup, n., clow, g.d., johnsen, s.j., hansen, a.w. & balling, n. 1998: past temperatures directly from the greenland ice sheet. science 282, 268–271. kaufman, d.s. et al. 2004: holocene thermal maximum in the western arctic (0–180°w). quaternary science reviews 23, 529–560. koç, n., jansen, e. & haflidason, h. 1993: palaeoceanographic reconstructions of surface ocean conditions in the greenland, iceland and norwegian seas through the last 14 ka based on diatoms. quaternary science reviews 12, 115–140. lamb, h.h. 1995: climate history and the modern world, 433 pp. london: routledge. mikkelsen, n., kuijpers, a. & arneborg, j. 2008: the norse in greenland and late holocene sea-level change. polar record 44, 45–50. ribergaard, m.h. 2007: oceanographic investigations off west greenland 2006. northwest atlantic fisheries organization, scientific council documents 07/1(n5339), 48 pp. seidenkrantz, m.-s., roncaglia, l., fischel, a., heilmann-clausen, c., kuijpers, a. & moros, m. 2008: variable north atlantic climate seesaw patterns documented by a late holocene marine record from disko bugt, west greenland. marine micropaleontology 68, 66–83. sparrenbom, c.j., bennike, o., björck, s. & lambeck, k. 2006: holocene relative sea-level changes in the qaqortoq area, southern greenland. boreas 35, 171–187. weidick, a., kelly, m. & bennike, o. 2004: late quaternary development of the southern sector of the greenland ice sheet, with particular reference to the qassimiut lobe. boreas 33, 284–299. 80 authors’ addresses n.n.p. & n.m., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen, denmark. email: nnp@geus.dk m.d.p., department of geography and geology, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. a.s.s., department of geology, university of tromsø, n-9037 tromsø, norway. rosa_2008:rosa-2008 01/07/09 15:49 side 80 geological survey of denmark and greenland bulletin 15, 2008, 17-20 more than 80% of the present-day oil and gas production in the danish part of the north sea is extracted from fields with chalk reservoirs of late cretaceous (maastrichtian) and early paleocene (danian) ages (fig. 1). seismic reflection and in version data play a fundamental role in mapping and characterisation of intra-chalk structures and reservoir properties of the chalk group in the north sea. the aim of seismic inversion is to transform seismic reflection data into quantitative rock properties such as acoustic impedance (ai) that provides information on reservoir properties enabling identification of porosity anomalies that may constitute potential reservoir compartments. petrophysical analyses of well log data have shown a relationship between ai and porosity. hence, ai variations can be transformed into porosity variations and used to support detailed interpretations of porous chalk units of possible reservoir quality. this paper presents an example of how the chalk team at the geological survey of denmark and greenland (geus) integrates geological, geophysical and petrophysical information, such as core data, well log data, seismic 3-d reflection and ai data, when assessing the hydrocarbon prospectivity of chalk fields. chalk chalk is a pelagic carbonate sediment, formed from settling of calcareous phytoplankton remains (i.e. coccoliths). the north sea chalk is practically a monomineralic carbonate reservoir rock that consists of 96–99% calcite (caco3), opal (radiolarians, diatoms and sponge spicules) and small amounts of clay minerals washed into the basin from land (håkans son et al. 1974; hancock 1975). as a reservoir rock, chalk is characterised by high porosity (25–50%) and low perme a b i lity (0.1–10 millidarcies, md; jørgensen & andersen 1991). de pending on permeability, chalk can be either a reservoir or a seal. due to the dominating calcitic nature of chalk, a robust empirical correlation exists between ai and total porosity (phit). hence, ai values can be transformed to phit and the relationship can be visualised by cross-plots of phit versus ai. variations in porosity obtained from inverted seismic data reflect changes in clay content or clay mineralogy, pore geometry, pore fluid and fracture characteristics (anderson 1999). using seismic attributes, it is also possible to distinguish between clean porous chalk with high-amplitude continuous reflectivity and clean tight chalk with low-amplitude discontinuous reflectivity. furthermore, the gamma-ray log – sonic-log correlation method can be used to distinguish between reservoir and non-reservoir intervals within the chalk group. the gamma-ray log can distinguish clean chalk intervals from more clay-rich chalk intervals, and the sonic log can be used as a porosity indicator. the gamma-ray log – sonic-log correlation method has been developed because gamma-ray and sonic logs are the most commonly available log types for north sea wells (britze et al. 2000). multi-disciplinary approach the kraka field an example from the kraka field in the southern part of the salt dome province in the danish north sea illustrates how a multidisciplinary approach can be used to identify and delineate porosity anomalies in the maastrichtian reservoir interval. a combination of detailed seismic interpretation, geophysical imaging of porosity variations in the danish north sea chalk tanni abramovitz © geus, 2008. geological survey of denmark and greenland bulletin 15, 17–20. available at: www.geus.dk/publications/bull 17 fig. 1. map of the central part of the north sea, showing the distribution of oiland gasfields. modified from vejbæk et al. (2007). 56°n oil in chalk gas in chalk field at other level 3°e 4°e 5°e 50 km uk norway denmark dan field kraka field germany the netherlands 500 km analysis of ai variations obtained from seismic inversion data, stratigraphic log correlation, petrophysical log analysis and rock physics analysis was used. the kraka field is a 60 km2 anticlinal structure with a four-way dip closure at top-chalk level. it has been induced through salt tectonics by up-doming of the chalk layers over a salt pillow (klinkby et al. 2005). the main chalk reservoir units are found in the danian ekofisk formation and the maastrichtian tor formation, which are characterised by high porosity (24–32%) on the crest of the structure and low matrix permeability of less than 1 md in the danian and 2–3 md in the maastrichtian units, and an effective permeability of 8–10 md due to fracturing. the oil zone is characterised by high water content (>50%) and limited thickness (70 m), and the gas cap is thin, less than 8 m. the free water level in the kraka field dips to the south-east according to thomasen & jacobsen (1994). geophysical interpretation integrating detailed seismic horizon interpretation and gamma-ray log/sonic-log correlation enables a consistent division of the chalk group and correlation of the chalk group units between wells in the study area. the interpretation of the gamma-ray log/sonic log of the anne-3 well in the kraka field (figs 1, 2) illustrates how reservoir and nonreservoir intervals can be identified. the maastrichtian units m1 to m4 were described by klinkby et al. (2005), who interpreted an intra-maastrichtian seismic horizon (base of unit m4; fig. 2) as the base of the reservoir interval in the kraka–dan area, separating porous chalk from tighter chalk below. as part of a major study of the internal chalk structures in the southern part of the danish north sea, several additional seismic horizons have been mapped including a porous unit (m5 on fig. 2) in the maastrichtian chalk reservoirs, which is present at a slightly deeper level than the previously interpreted base of the reservoir interval. the base of the m5 unit is seen as a distinct, relatively highamplitude continuous reflection (trough) in the seismic data. the horizon can be correlated from the kraka wells to olga1x and the dan field wells (m-1x, m-2x, m-8x, m-9x and m-10x), where it corresponds to a consistent regional reflection between two intra-maastrichtian intervals. the reflector marking the base of the m5 unit changes character and becomes weaker towards the north-east in the direction of the alma-1x and alma-2x wells. near the anne-3 well and farther to the south-west, it merges with a deeper lying intrachalk seismic horizon (near top hod; fig. 3). here, an e–w-striking ai profile crossing the anne-3 well shows pronounced lateral variability in ai values within the m5 unit. on the eastern flank of the kraka structure, the m5 unit is dominated by low ai values, corresponding to higher porosities. in general, to the north and to the west the base of the m5 unit separates the porous maastrichtian units with lower ai values from the underlying units with higher ai values (figs 3, 4). 18 sonic (dt) 140 µs/ft 50 sw 0 1 phie 75 0% 5 api 40 gamma ray 5 api 40 gamma ray d an ia n po ro us d an ia n tig ht m aa st ri ch tia n po ro us m1 m2 m3 m4 m5 tr ue v er tic al d ep th ( m b el ow s ea le ve l) gr/phie log 1800 1850 1900 1950 2000 top chalk intra-danian top maastrichtian (top tor) intra-maastrichtian (base m4) intra-maastrichtian (base m5) near top hod porosity in clean chalk >33% 25–33% <25% intra-chalk unconformity fig. 2. gr-dt log for the anne-3 well, kraka field. gamma-ray, sonic, effective porosity and water saturation log curves are shown for the different reservoir units of the danian –maastrichtian reservoir interval. high gr chalk is shown in brown. dt, sonic; gr, gamma ray; sw, total water saturation; phie, effective porosity. modified from klinkby et al. (2005). transforming acoustic impedance to porosity based on petrophysical log-porosity evaluations of 14 wells from the southern part of the danish central graben, a robust relationship between log-derived ai and phit has been established for the chalk group (kristensen & ander sen 2008). ai can be converted to phit using a second order polynomial regression. the following relationship was found for clean chalk (<2% clay) where phit is close to the effective porosity (phie): phit = 0.729 – 7.08 × 10–5ai + 1.55 × 10–9ai2 where porosity is given in fraction and ai in g cm–3 × msec–1. in order to display the ai variations of the m5 unit on a map, the lowest ai values were extracted from the inverted seismic data within the target interval. the ai map of the m5 unit shows the presence of an area with low ai values on the south-eastern flank of the kraka structure (fig. 5). the ai–phit relationship allowed the generation of a total porosity map (fig. 6) for the target interval using the seismically derived ai values (fig. 5). the porosity map shows an area with higher porosities (30–35%) east of the anne-3 well on the south-eastern flank of the kraka structure. this indicates the existence of high porosity values in the deeper part of the m5 unit. in addition, the seismically derived phit values show good agreement with the gr-dt log values estimated for the m5 unit (fig. 2). the area with low ai values (labelled a on the map; fig. 5) corresponds to label a on the north –south orientated, inverted seismic profile (fig. 4) and to the high porosity area labelled a’ on the porosity map (fig. 6). the area labelled b on the n–s-orientated inverted, seismic profile (fig. 4) illustrates the abrupt transition from higher to lower ai values within the mapped unit (equal to b on fig. 6). this corresponds to an area with relatively low porosity marked by b’ on fig. 6. similarly, the lower porosities (<25%) labelled c’ on fig. 6 and located in the saddle between the kraka and dan structures correspond to an area with high ai values, labelled c on figs 4, 5. conclusions and outlook a multi-disciplinary approach is a powerful method for detecting and predicting intervals of high porosity of chalk in the north sea. this study indicates the existence of high porosity values at a slightly deeper level than the previously interpreted base of the reservoir interval on the south-eastern flank of the kraka field. a porosity map based on inverted seismic data may contribute as input to a 3-d reservoir model for further calculation of hydrocarbon volumes at different reservoir property and fwl scenarios. 19 tw o -w ay t ra ve l t im e (m se c) 2000 1900 2100 2500 2300 w e 1800 2400 2200 top chalk intra-danian top maastrichtian (tor) intra-maastrichtian (base m4) intra-maastrichtian (base m5) base chalk near top hod a-7ca-7c interval of interestinterval of interestinterval of interest 1 km anne-3 10 000 9000 8000 7000 acoustic impedance (103 kg m–2 s–1) a-7c fig. 3. an e–w-orientated acoustic impedance profile extracted from the inverted seismic data of the merged 3-d survey (produced for mærsk oil and gas as) crossing the anne-3 well, kraka field. n s 2000 1900 2100 2300 tw ow ay t ra ve l t im e (m se c) 1800 2400 2200 top chalk intra-danian top maastrichtian (tor) intra-maastrichtian (base m4) intra-maastrichtian (base m5) base chalk near top hod interval of interestinterval of interestinterval of interest 1 km a-7c acoustic impedance (103 kg m–2 s–1) 10 000 9000 8000 7000 c b a c b a fig. 4. a n–s-orientated acoustic impedance section from the dan to kraka fields. labels a, b and c are discussed in the text. references anderson, j.k. 1999: the capabilities and challenges of the seismic method in chalk exploration. in: fleet, a.j. & boldy, s.a.r. (eds): petroleum geology of north-west europe. proceedings of the 5th petroleum geology conference, 939–947. london: geological society. britze, p., nielsen, e.b., dahl, n. & haug, s. 2000: north sea chalk porosity resolved by integration of seismic reflectivity and well log data. abstract p100, 64th eage conference & exhibition, florence. abstract p100, 4 pp. håkansson, e., bromley, e.g. & perch-nielsen, k. 1974: maastrichtian chalk from north west europe – a pelagic shelf sediment. special publication, international association of sedimentologists 1, 211–233. hancock, j.m. 1975: the petrology of the chalk. proceedings of the geologists’ association 86, 499–535. jørgensen, l.n. & andersen, p.m. 1991: integrated study of the kraka field. society of petroleum engineers paper, 14 pp. (http://dx.doi.org/ 10.2118/23082-ms). klinkby, l., kristensen, l., nielsen, e.b., zinck-jørgensen, k. & stem merik, l. 2005: mapping and characterisation of thin chalk reservoirs using data integration: the kraka field, danish north sea. petroleum geoscience 11, 113–124. kristensen, l. & andersen, c. 2008: log-derived acoustic impedance versus porosity and porosity versus depth trends in the chalk group: examples from the southern danish central graben. danmarks og grønlands geologiske undersøgelse rapport 2008/13, 33 pp. thomasen, j.b. & jacobsen, n.l. 1994: dipping fluid contacts in the kraka field, danish north sea. society of petroleum engineers paper, 10 pp. (http://dx.doi.org/10.2118/28435-ms). vejbæk, o.v., bidstrup, t., britze, p., erlström, m., rasmussen, e.s. & sivhed, u. 2007: chalk depth structure maps, central to eastern north sea, denmark. geological survey of denmark and greenland bulletin 13, 9–12. 20 author’s address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: tab@geus.dk 7000 8000 9000 2 km 10000 acoustic impedance (103 kg m–2 s–1) 55°30′ 55°30′ 55°25′ m-1x m-9x a-10c a-4h a-4p a-2x a-1x a-6a a-6ll a-9 a-7ca-6l anne-3 a-5p a-10p m-8x m-2x m-10x mff-19p c b a 5°00′ 5°00′ 5°05′ 5°10′ a-8 olga-1x 5°05′ 5°10′ 20 25 30 35 phit (%) 55°25′ 55°30′ 55°30′ 55°25′ m-1x m-9x a-10c a-4h a-4p a-2x a-1x a-6a a-6ll a-9 a-7ca-6l anne-3 a-5p a-10p m-8x m-2x m-10x mff-19p c’ b’ a’ 5°00′ 5°00′ 5°05′ 5°10′ 5°05′ 5°10′ a-8 olga-1x 2 km fig. 5. map showing the distribution of the lowest acoustic impedance values in unit m5 in the area of the kraka and dan fields. labels a, b and c correspond to the labels on fig. 4. the black stippled lines show the position of the two profiles shown in figs 3 and 4. for location see fig. 1. fig. 6. porosity map for the m5 unit in the area of the kraka and dan fields derived from the ai-phit transform. porosity anomalies labelled a’, b’ and c’ correspond to a, b and c on figs 4 and 5 as discussed in the text. the black stippled lines show the position of the two profiles shown in figs 3 and 4. for location see fig. 1. research article smith & moseley 2022: geus bulletin 49. 8298. https://doi.org/10.34194/geusb.v49.8298 1 of 23 the karst and palaeokarst of north and north-east greenland – physical records of cryptic geological intervals m. paul smith*1  , gina e. moseley2  1oxford university museum of natural history, oxford, uk; 2institute of geology, university of innsbruck, innsbruck, austria abstract carbonate rocks of neoproterozoic to silurian age are abundantly distributed around the coasts of north and north-east greenland. palaeokarst horizons are particularly well developed within the portfjeld formation (ediacaran – earliest cambrian) and beneath the buen formation (cambrian series 2), and there are caves within ordovician limestones infilled by caledonian molasse of middle devonian age. the youngest karst is a series of caves distributed from hall land in western north greenland to kronprins christian land in eastern north greenland. caves within ordovician carbonates in freuchen land are currently the northernmost documented karst caves globally. the caves are mainly open phreatic conduits, any fill that is present is unlithified, and cave collapse is limited to minor breakdown associated with frost shattering. these geologically young caves are consistently located up to a few 100 m beneath the distinctive plateau that characterises the topography of the northern coast, and their identical context suggests that they developed in a single phase of speleogenesis. the caves are exposed where the plateau has been incised by outlet glaciers from the greenland ice sheet. the timing of cave development in north greenland is constrained by the midto late-miocene (15–5 ma) uplift of the plateau surface and the onset of fjord-forming glaciation in the latest pliocene – earliest pleistocene (c. 2.7–2.5 ma). the evidence suggests that phreatic caves in the southern part of northeast greenland, on c. h. ostenfeld nunatak, are of a broadly similar age. the caves of north and north-east greenland offer a glimpse of large-scale phreatic drainage systems that developed below an uplifted coastal peneplain during neogene time. they preserve an important part of the geological history of north and north-east greenland that is otherwise absent from the physical geological record. introduction carbonate rocks of neoproterozoic–silurian age extend for 1000 km from west to east across north greenland and then intermittently down the 1300 km length of the n–s-oriented east greenland caledonides (figs 1, 2, 3). the karst developed in these rocks includes both geologically young caves truncated by later glaciation together with regionally extensive palaeokarst horizons developed in rocks of the latest ediacaran to ordovician age (fig. 4), which were generated by a combination of relative sea-level lowstands, periods of subaerial erosion at sequence boundaries, orogenic topography and passive margin uplift. *correspondence: paul.smith@oum.ox.ac.uk received: 23 oct 2021 accepted: 10 mar 2022 published: 11 may 2022 keywords: caves, cenozoic, karst, neogene, miocene, palaeokarst, paleokarst abbreviations: afta: apatite fission track analysis a.s.l.: above sea level cai: colour alteration index dem: digital elevation model gd: grottedal (used in alphanumeric scheme of moseley 2016) geus: geological survey of denmark and greenland ggu: geological survey of greenland ics: international commission on stratigraphy iugs: international union of geological science ups: upper planation surface lps: lower planation surface geus bulletin is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: jon r. ineson (geus, denmark) reviewed by: jan tveranger (norce norwegian research centre, norway), peter frykman (geus, denmark) funding: see page 21 competing interests: see page 21 additional files: see page 21 https://doi.org/10.34194/geusb.v49.8298 https://orcid.org/0000-0002-5141-1577 https://orcid.org/0000-0002-5618-5759 mailto:paul.smith@oum.ox.ac.uk smith & moseley 2022: geus bulletin 49. 8298. https://doi.org/10.34194/geusb.v49.8298 2 of 23 www.geusbul let in.org from the early cambrian to the silurian (521–420 ma), greenland lay astride the palaeo-equator as part of laurentia (mac niocaill & smethurst 1994), and north greenland is one of the few places on the laurentian margin where an ancient shelf–slope break is preserved with minimal tectonic overprint. the north-east corner of greenland, where the east and north coast meet, represents an original promontory of the laurentian continent (derby et al. 2012, fig. 1) and, in consequence, the cambrian– silurian successions on these two margins have contrasting stratigraphic frameworks and tectonic histories. the cambrian–ordovician succession of the east coast comprises an initial siliciclastic interval of c. 200 m thickness overlain by 2.7 km of predominantly subtidal carbonates that span only 60 million years. this succession was once contiguous with equivalent sectors that lay farther to the south on the laurentian margin, including north-western scotland and western newfoundland (swett & smit 1972; swett 1981; smith & rasmussen 2008; raine & smith 2012). in contrast, the cambrian–ordovician on the north coast is stratigraphically thinner, with a substantial hiatus in eastern north greenland caused by uplift and erosion on the apex of the laurentian promontory (peel & smith 1988; smith 2000). depositional continuity and stratigraphic correlation extend westwards across nares strait into the coeval sedimentary successions of nunavut, canada (trettin 1991). an extended period of sea-level lowstands during the neoproterozoic led to the extensive peneplanation of continental interiors (peters & gaines 2012). late cambrian to ordovician sea-level highstands constituted the highest absolute sea levels of the phanerozoic (miller et al. 2005), which in combination with low relief in laurentia led to carbonate deposition extending almost continuously for 6500 km from modern new mexico to greenland, and the development of a long-lived, non-uniformitarian geological and geomorphological feature – the great american carbonate bank (derby et al. 2012). for the most part, it is the carbonate deposited as part of the great american carbonate bank in which the karst and palaeokarst of greenland are developed. the aim of this paper is to provide a complete documentation of the palaeokarst and younger caves recorded to date in north and north-east greenland and to ascertain the timing of their formation, particularly the temporal constraints on the speleogenesis of the younger caves. the geographical divisions used in this study follow the conventions of the geological survey of denmark and greenland (geus) as outlined by dawes et al. (2016), wulff land kronprins christian land nyeboe land nares land peary land 60°w 40°w 20°w 10°w50°w 80°n km fig. 2 morris bugt & washington land groups (upper ordovician–silurian) palaeokarst localities cave localities sg nof naf hg lincoln hav wandel hav fig. 3 warming land 1000 nb victoria fj. ind ependence fj. johannes v. jensen land nioghalvfjerdsfjorden hall land danm ark fjord se fl hagen fjord portfjeld formation (ediacaran–cambrian) øm sp fig. 1 map of north greenland showing the outcrop of sandbian (upper ordovician) to wenlock (silurian) shelf limestones. palaeokarst localities are indicated by a red dot and cave localities by a black dot; dots may represent more than one instance where localities are closely spaced. for full details of all known caves, see supplementary information. fl: freuchen land. hg: hjørnegletscher. nb: newman bugt. naf: navarana fjord. nof: nordenskiöld fjord. øm: øvre midsommersø; se: sæfaxi elv. sp: sirius passet. sg: steensby gletscher. linework based on the 1:500 000 ggu and geus geological maps (bengaard & henriksen 1986; henriksen 1989; jepsen 2000). https://doi.org/10.34194/geusb.v49.8298 http://www.geusbulletin.org smith & moseley 2022: geus bulletin 49. 8298. https://doi.org/10.34194/geusb.v49.8298 3 of 23 www.geusbul let in.org with the boundary between eastern north and northeast greenland placed at nioghalvfjerdsfjorden; the boundary between eastern and central north greenland at hagen fjord; and that between central and western north greenland at victoria fjord (fig. 1). palaeokarst in greenland limestones and other carbonate rocks occur commonly in the neoproterozoic–silurian shelf succession of the laurentian margin in north greenland (figs 1, 2) and in the cambrian–ordovician of the north-east greenland coast (fig. 3). this widespread occurrence has provided the potential for the development of karstic horizons at multiple points in the geological history of the area in response to relative changes in sea level. there are three particularly well-developed palaeokarst horizons with regionally extensive distributions. intra-ediacaran palaeokarst the portfjeld formation of north greenland extends from the inner shelf of the franklinian basin, in a southern outcrop belt to the west and east of øvre midsommersø (øm in fig. 1; 82.23°n, 36.10°w), and a northern outcrop belt along the line of the cambrian shelf edge, including sirius passet (sp in fig. 1; higgins et al. 1991; willman et al. 2020). the 200–700 m thick formation comprises pale grey dolostones with some marly limestones and sandstones, but the southern outcrop belt contains a major unconformity that separates midto late-ediacaran (<570 ma) strata from a younger succession that is of the latest ediacaran or earliest cambrian age (fig. 4; ineson & peel 2011; willman et al. 2020). the dating of the older unit is confirmed by the presence of the shuram-wonoka ∂13c anomaly (570–560 ma; willman et al. 2020). 20°w 80°n km 20°w80°n hg ingolf fjord sæfaxi elv marmorvigen vandredalen grottedal centru msø da nm ar k fj or d hekla sund dijm ph na sun d kh gr æ se lv fs 200 prinsesse caroline-mathilde alper sk jo ld un ge el v morris bugt & washington land groups (upper ordovician–silurian) wandel valley formation (lower–middle ordovician) kap holbæk formation (cambrian series 2) proterozoic sedimentary and igneous rocks palaeokarst localities cave localities faultthrust fig. 2 geological map of kronprins christian land, north-east greenland, showing the distribution of sandbian (upper ordovician) to wenlock (silurian) shelf limestones in which caves are developed (black dots). a regionally extensive palaeokarst horizon is preserved where cambrian series 2 quartz arenites of the kap holbæk formation rest on carbonate units such as the tonian (lower neoproterozoic) fyns sø formation (red dots). fs: fyn sø. hg: hjørnegletscher. kh: kap holbæk. linework based on primary field mapping in 1994, 1995, and 2019, together with the 1:500 000 geus geological map (jepsen 2000). https://doi.org/10.34194/geusb.v49.8298 http://www.geusbulletin.org smith & moseley 2022: geus bulletin 49. 8298. https://doi.org/10.34194/geusb.v49.8298 4 of 23 www.geusbul let in.org at the western end of øvre midsommersø (fig. 1), willman et al. (2020) also recorded well-developed palaeokarst at the unconformity surface within the portfjeld formation. the karst includes extensive intrastratal solution cavities and brecciation that extend for up to 40 m downwards from the subaerial surface represented by the unconformity. cavities are infilled by sandy breccias and cements. sub-cambrian series 2 palaeokarst the most regionally extensive palaeokarst surface is present beneath transgressive siliciclastic shelf sediments of cambrian series 2 age, assigned to the buen formation from warming land to danmark fjord and to the kap holbæk formation in kronprins christian land (figs 1, 2, 4; smith et al. 2004). in the north and west, the buen formation overlies carbonates of the portfjeld formation, fig. 3 geological map of albert heim bjerge and c. h. ostenfeld nunatak showing cave (black dots) and palaeokarst (red dot) localities in limestones of the cape weber formation (lower ordovician) and the heimbjerge formation (middle to upper ordovician). the ordovician rocks of this area are unconformably overlain by middle to late devonian molasse deposits of the vilddal and kap kolthoff groups. location of the map indicated on fig. 1. linework is based on primary field mapping in 1998 and the 1:500 000 geus geological map (escher 2001). wordie gletscher promenadedal vibeke sø vibeke elv godthåb golf steno land c. h. ostenfeld nunatak albert heim bjerge 74.3°n 74.2°n 74.1°n 74.1°n 74.2°n 74.3°n 23.5°w 23.0°w 22.5°w 23.5°w 23.0°w 22.5°w km g rantagletscher grantafjord 0 10 narwhale sound & heimbjerge formations (m–u ordovician) antiklinalbugt & cape weber formations (lower ordovician) undifferentiated quaternary proterozoic sedimentary and metasedimentary rocks palaeokarst localities cave localities fault cambrian units vilddal & kap kolthoff groups (middle–upper devonian) https://doi.org/10.34194/geusb.v49.8298 http://www.geusbulletin.org smith & moseley 2022: geus bulletin 49. 8298. https://doi.org/10.34194/geusb.v49.8298 5 of 23 www.geusbul let in.org whereas in kronprins christian land, the kap holbæk formation rests on older, tonian (early neoproterozoic) carbonates of the fyns sø formation (fig. 4). in places, the kap holbæk formation is preserved only within palaeokarst cavities of the older unit (smith et al. 1999). peary land – 82.2–82.8°n a well-preserved karstic surface exhibiting both exoand endokarst is seen at the top of the portfjeld formation at the sirius passet lagerstätte locality adjacent to j.p. koch fjord (sp in fig. 1; 82.79695°n, 42.21585°w), where it is ma 420 430 440 450 460 470 480 490 500 510 520 530 541 sequenceseries stagesystem nyeboe land – freuchen land kronprins christian land northeast greenland ediacaran devonian pridoli ludlow wenlock yrevo dnall upper middle lower naignoruf m ia ol in gi an 2 seire s te rr en eu vi an for stage 2 stage 3 stage 4 wul dru guz pai jia stage 10 tre flo dap dar san kat hir rhu aer tel s ilu ria n o rd ov ic ia n c am b ria n sauk i sauk ii sauk iii sauk iv i eonace p pit iii eonace p pit tip ii peary land group hb djævlekløften aleqatsiaq fjord kap jackson ryder gletscher group brønlund fjord group buen cape calhoun peary land groupsh odins fjord turesø børglum river kap holbæk heimbjerge narwhale sound cape weber antiklinalbugt dolomite point ella island bastion hyolithus creek kløftelv sjælland fjelde wandel valley portfjeld ii portfjeld i ph cape webster palaeokarst localities cave localities mixed clastic marine rocks peritidal dolostones subtidal/deep shelf limestones marine sandstones fig. 4 correlation chart of cambrian–silurian shelf units in central and western north greenland (nyeboe land – freuchen land), eastern north greenland (kronprins christian land) and north-east greenland (c. h. ostenfeld nunatak) showing the stratigraphic distribution of caves and palaeokarst. stratigraphic units are of formation rank unless otherwise stated. the left-hand columns follow the international union of geological science/international commission on stratigraphy (iugs/ics) standard (cohen et al. 2013); boundaries of megasequences and supersequences are from palmer (1981), golonka & kiessling (2002) and smith & rasmussen (2008). data for the greenland sections were compiled from smith (1985, 1991), tull (1988), smith & bjerreskov (1994), and huselbee (1998). for: fortunian. wul: wuliuan. dru: drumian. guz: guzhangian. pai: paibian. jia: jiangshanian. tre: tremadocian. flo: floian. dap: dapingian. dar: darriwilian. san: sandbian. kat: katian. hir: hirnantian. rhu: rhuddanian. aer: aeronian. tel: telychian. stratigraphic units: hb: hauge bjerge formation. ph: petermann halvø formation. sh: samuelsen høj formation. https://doi.org/10.34194/geusb.v49.8298 http://www.geusbulletin.org smith & moseley 2022: geus bulletin 49. 8298. https://doi.org/10.34194/geusb.v49.8298 6 of 23 www.geusbul let in.org disconformably overlain by quartz arenites of the basal buen formation (ineson & peel 2011; harper et al. 2019, fig. 2). the upper part of the portfjeld carbonates contains vertical and inclined karstic grikes and vadose fissures (kluftkarren) that are 0.1–3 m wide, together with shallow, infilled phreatic tubes immediately beneath the disconformity surface. all of the karstic cavities are infilled with fineto medium-grained quartz arenites that have millet-seed texture or with black muddy siltstone (fig. 5). in places a sequential fill of quartz arenites followed by muddy siltstone is observed, and the same depositional sequence is present in the most basal part of the buen formation immediately above the disconformity (harper et al. 2019, fig. 2). the same karstic surface is seen at the portfjeld–buen formation boundary at øvre midsommersø (øm in fig. 1), 100 km to the south-east of the sirius passet lagerstätte, where the most inboard shelf sediments of this age are seen in north greenland. at this locality, the palaeokarst is characterised by collapse dolines and other karstic collapse structures containing sandstone infills. kronprins christian land – 80–81°n although the sub-buen formation disconformity in peary land has well-preserved palaeokarst features and relief, the equivalent surface in kronprins christian land to the south-east (fig. 1) is notable for containing infilled cave systems of substantial size. in the west, the buen-equivalent kap holbæk formation rests unconformably on the fyns sø formation of the hagen fjord group (smith et al. 2004) along danmark fjord and skjoldungeelv to the south and in the east along the inner part of ingolf fjord and at marmorvigen in sæfaxi elv (se in figs 1, 2). the best-developed sub-cambrian palaeokarst is at hjørnegletscher (hg in figs 1, 2; 80.66891°n, 19.49717°w), which enters ingolf fjord from the north. early cambrian sandstones of the kap holbæk formation rest unconformably on tonian (early neoproterozoic) dolostones of the fyns sø formation, and the succession is strongly deformed by west-vergent folds developed in the footwall of the vandredalen thrust (smith et al. 1999; higgins et al. 2004). despite this deformation, a well-preserved, multi-phase palaeokarst system is preserved in the uppermost part of the fyns sø formation (smith et al. 1999). the unconformity surface has two deep channels that are, respectively, 40 m wide by 7 m deep and 30 m wide by 9 m deep (fig. 6a). one of the channels has a bedding plane-controlled solution cavity that extends beyond the channel margin (fig. 6a, arrowed). the lowest sediment filling the channel is a moderately sorted, fineto coarse-grained black sandstone and is overlain by a well sorted, very fine sandstone and in turn by a fineto medium-grained quartz arenite. a large number of phreatic conduits are present beneath the unconformity surface, with tubes that range from less than 1 m across up to one that is 15 m wide and at least 5 m high (figs 6a, b). some of the conduits are circular in cross-section, but the majority extend laterally along bedding planes. the sedimentary fill and the internal stratigraphy are identical to that of the large channels (smith et al. 1999). the caves extend for only 12 m below the preserved unconformity surface and are perched on an aquiclude of insoluble black siltstones, which probably inhibited their downward development (smith et al. 1999). it is evident from the relative positions of the surface channels and the phreatic tubes that multiple stages of karst development occurred. the first phase was the development of phreatic tubes at an unknown depth beneath the unconformity, as the amount of erosion at the unconformity surface is a b fig. 5 sub-cambrian palaeokarst developed within the upper 2 m of the portfjeld formation (ediacaran or lowermost cambrian) at the sirius passet lagerstätte locality, north-western peary land (fig. 1; harper et al. 2019) and infilled with clastic sediments of the buen formation (cambrian, series 2). a: steeply inclined fissure infilled by millet-seed quartz arenite, margins arrowed, with a fin of dolostone dividing the fissure into two close to the disconformity surface (skyline of block). b: inclined fissure (arrow) and sub-horizontal cavity (above, in shadow) infilled with metamorphosed muddy siltstones. skyline of foreground is again the unconformity surface. yellow rule is 1 m. https://doi.org/10.34194/geusb.v49.8298 http://www.geusbulletin.org smith & moseley 2022: geus bulletin 49. 8298. https://doi.org/10.34194/geusb.v49.8298 7 of 23 www.geusbul let in.org unconstrained. following relative base-level fall, incision of the surface channels occurred, and these must have post-dated cave formation as they extend to a greater depth below the unconformity surface than some of the phreatic caves (fig. 6a). the lateral, bedding plane-controlled solution cavities adjacent to the channels are best interpreted as truncated phreas (sub-water table cave conduit) since they have a similar morphology to the adjacent caves (fig. 6a, arrowed). the morphology of this phreas may have been modified by caledonian thrusting associated with the nearby vandredalen thrust, leading to flattening and an accentuated low profile, but they must originally have been bedding plane-controlled conduits. the final phase of development was infill by the lower cambrian sediments of the kap holbæk formation. although there is a substantial age difference between the host fyns sø formation carbonates and the kap holbæk formation infill (>200 million years), the formation of the phreatic tubes and the incision of the channels were probably cambrian in age, evidenced by the lack of collapse in the caves and the presence of identical sedimentary fills in the caves and channels elsewhere above the unconformity. in sæfaxi elv (figs 1, 2; 80.08385°n, 20.42261°w), the tonian dolostones of the fyns sø formation are overlain by a 1 m thick coarse sandstone with imbricated pebbles of dolostone and quartz arenite (smith et al. 1999), and this is in turn overlain by floian (lower ordovician) dolostones of the wandel valley formation (fig. 6e). lithologically, this sandstone is best interpreted as the lowest bed of the ordovician succession. however, sandstone-filled fissures 20–60 cm wide and up to 2 m long (fig. 6d) extend downwards from the sub-planar unconformity and at a depth of around 10 m connect with a network of subhorizontal phreatic conduits (smith et al. 1999, figs 6, 7). these tubes are circular to elliptical and range in width from a few decimetres up to 3 m. one distinctive example is an hour-glass shaped tube, 1.9 m high and up to 0.8 m wide (fig. 6c). the fissures and conduits are infilled by a distinctively golden-brown weathering quartz arenite of coarse sand – granule grade, which contrasts with the finer sandstone bed at the base of the wandel valley formation. this lithological contrast and the presence of quartz arenite intraclasts in the boundary sandstone bed suggests that the palaeokarst fill is an older unit, and fränkl (1955) suggested it was the kap holbæk formation. the similarity to a b c d e wv sst fs kh fs kh fs kh phreatic conduit channel 1 channel 2 solution cavity infilled cave fig. 6 sub-cambrian palaeokarst developed within tonian (neoproterozoic) carbonate rocks of the fyns sø formation (fs) of kronprins christian land and infilled with quartz arenites of the kap holbæk formation (kh; series 2, cambrian). a: channel and phreatic tube complex developed in the upper 12 m of the fyns sø formation on the west side of hjørnegletscher (figs 1, 2). this is the same conduit as in b, where there is a person for scale. b: close-up of the phreatic conduit seen on the left side of a. c: sub-horizontal phreatic tube with hourglass profile on the south side of sæfaxi elv at marmorvigen (figs 1, 2), located 10 m below the unconformity surface and infilled with very coarse sand to granule-grade quartz arenite. d: vertical vadose fissure in subhorizontal dolostones and infilled with coarse-grained quartz arenite. e: dark-coloured sandstones (sst) overlying the fyns sø formation and in turn overlain by floian (lower ordovician) pale-coloured dolostones of the wandel valley formation (wv). divisions on the survey pole are 20 cm; hammer is 28 cm long. https://doi.org/10.34194/geusb.v49.8298 http://www.geusbulletin.org smith & moseley 2022: geus bulletin 49. 8298. https://doi.org/10.34194/geusb.v49.8298 8 of 23 www.geusbul let in.org the kap holbæk formation and cave infills at ingolf fjord, 65 km to the north, strongly supports this correlation. it is noteworthy that the sæfaxi elv caves are also developed a very short distance below the unconformity surface, indicating a very shallow vadose zone, which may also be true of the hjørnegletscher locality, where karst formation was constrained by the aquiclude. the sæfaxi elv and ingolf fjord localities both lie in the deformed footwall of the vandredalen thrust, within the caledonian thrust belt (higgins et al. 2004), but additional sub-cambrian palaeokarst localities are present in the foreland to the west, along danmark fjord and to the south of it. to the east of fyn sø (fs in fig. 2), the uppermost levels of the fyns sø formation contain fissures and irregular tubes 10–50 cm in diameter infilled with iron-stained, mediumto coarse-grained quartz arenites (smith et al. 1999). the fyns sø formation is here overlain by the kap holbæk formation, and the unconformity surface has a karstic erosion surface with up to 3 m of relief. farther north along the west side of danmark fjord, quartz arenites again infill irregular fissures up to 2 m deep and 0.2 m wide in the uppermost fyns sø formation (smith et al. 1999). the distribution of sub-cambrian series 2 palaeokarst wherever the buen and kap holbæk formations unconformably overlie carbonate units indicates that this is a major geomorphological surface with exoand endokarst development of regional extent. that region is bounded by sirius passet in the northwest, øvre midsommersø in the south-west, north-eastern peary land in the north-east and sæfaxi elv in the south-east (fig. 1), equating to a minimum area of over 50 000 km2. hypogene karst in kronprins christian land moseley (2016, p. 60) identified one probable example of hypogene karst formed by hot fluids in grottedal, kronprins christian land, on the basis of its morphology (fig.  2; 80.37369°n, 21.72854°w). hypogene cave (gdx in the alphanumeric scheme of moseley 2016 where gd refers to grottedal) comprises a horizontal tube of 30–35 cm diameter that can be traced for around 3 m in limestones of the odins fjord formation (llandovery, silurian). supporting evidence for a high-temperature origin comes from the sequential fill of dense, laminated calcite followed by coarse-grained, euhedral calcite that contrasts markedly with other cave fills in the region. remaining void is infilled by buff-coloured silt. the conduit is probably of caledonian origin as this is the only post-depositional event in the grottedal area likely to have generated sufficiently high temperatures. rasmussen & smith (2001) obtained conodonts with a colour alteration index of cai 4 in grottedal, corresponding to burial temperatures of 190– 300°c (epstein et al. 1977) and in turn indicating a minimum burial depth of 6.8 km beneath the overburden of caledonian thrust sheets and foreland basin deposits (the latter now removed by erosion; rasmussen & smith 2001). sub-devonian palaeokarst – wordie gletscher, east greenland, 75°n in the north-east greenland caledonides, minor palaeokarst features are seen wherever devonian conglomerates of the post-orogenic molasse (vilddal group; givetian, middle devonian) overlie carbonate units within the orogen (fig. 3). palaeokarst is best developed in the relatively pure subtidal limestones of the ordovician cape weber and heimbjerge formations (figs 3, 4). on albert heim bjerge (fig. 3; 74.1076°n, 22.9273°w) at c. 100 m above sea level (a.s.l.), two caves infilled with coarse breccias of the solstrand formation (vilddal group) were identified during the regional mapping programme in 1998 at the eastern end of albert heim bjerge, adjacent to wordie gletscher (fig. 3). the sub-quadrate phreatic conduits are 40 cm wide (fig. 7a) and 60 cm wide a b fig. 7 examples of quadrate phreatic conduits (a, b) developed within subtidal limestones of the cape weber formation (tremadocian–floian, lower ordovician) infilled with devonian conglomerates of the vilddal group (middle devonian) that mainly comprise blocks of red-stained ordovician carbonates in a clastic matrix; the eastern end of albert heim bjerge adjacent to wordie gletscher (fig. 3). divisions on the survey pole are 20 cm, and lens cap is 7 cm. https://doi.org/10.34194/geusb.v49.8298 http://www.geusbulletin.org smith & moseley 2022: geus bulletin 49. 8298. https://doi.org/10.34194/geusb.v49.8298 9 of 23 www.geusbul let in.org (fig. 7b), respectively. they are developed within subtidal limestones of the cape weber formation (tremadocian– floian, lower ordovician), around 75 m above the base of the unit, and the infill of solstrand formation comprises blocks of red-stained ordovician carbonates in a finer clastic matrix. the sub-devonian unconformity surface is deeply erosive into folded cambrian–ordovician strata, such that on albert heim bjerge the devonian overlies floian cape weber formation and younger carbonates, but on c. h. ostenfeld nunatak, 20 km to the north, it rests on darriwilian limestones of the heimbjerge formation (fig. 4). the unconformity surface is highly irregular at a smaller scale and has a variety of karstic pockets infilled with devonian sediment. in places, the c. 70 m thick devonian conglomerate overlying the cape weber formation at this locality has such a high proportion of carbonate clasts that it exhibits karstic weathering itself. the stratigraphic and tectonic context (higgins et al. 2008) constrains phreatic cave formation on albert heim bjerge to the interval between 460 ma (the youngest pre-orogenic sedimentary unit; smith & rasmussen 2008) and 388 ma (the givetian infill). caves of north and north-east greenland the younger caves of greenland are of neogene age, in contrast to the older palaeokarst horizons, and lack lithified, pre-cenozoic fills. they characteristically occur high in fjord walls and other cliffs, where they have been truncated and exposed by glaciation. in north greenland, they are characteristically located a few tens to hundreds of metres below the distinctive plateau surface that is developed from hall land eastwards to kronprins christian land (figs 1, 8). the caves are predominantly former phreatic conduits, many of which are very large – up to tens of metres in width. there appears to be strong lithological/stratigraphic control over the position and perhaps, in turn, speleogenesis of many of the caves. all of those explored to date are choked by ice or unlithified sediment fill within a few metres or tens of metres (moseley 2016; moseley et al. 2020). nevertheless, the caves do have occurrences of speleothems (secondary mineral deposits precipitated in caves) that are beginning to yield important information about pleistocene climate in north greenland (moseley et al. 2021). in addition, the geological location and context of the caves in north and north-east greenland provide an insight into the landscape evolution of these regions, in time intervals that are otherwise unrepresented by a stratigraphic record. wordie gletscher, east greenland – 74°n the southernmost documented caves in greenland without palaeozoic fill are located in ordovician carbonates on albert heim bjerge and c. h. ostenfeld nunatak 60°w 50°w 82°n nares land wulff land nyeboe land hall land freuchen land warming land victoria fjord j. p. koch fjord 17 00 m wul8 wul6 nf1 fl1 nl2 nl11 wal 9 wal8 wal6 0 1700altitude (m a.s.l.) perm in land pl1 fig. 8 digital elevation model (dem) of central and western north greenland showing the caves described and figured in the text (data from arctic dem, porter et al. 2018). the caves are developed immediately beneath the plateau surface, at an altitude of 800–1000 m, that is seen in the southern parts of the area, from freuchen land westwards to hall land. the dark blue peaks are ice caps sitting on the plateau surface. documentation and illustration of other caves discovered during aerial reconnaissance are provided in supplementary file s1. the colour ramp extends from sea level to 1900 m a.s.l., and the shift to the darkest blue tone is at 1700 m a.s.l. https://doi.org/10.34194/geusb.v49.8298 http://www.geusbulletin.org smith & moseley 2022: geus bulletin 49. 8298. https://doi.org/10.34194/geusb.v49.8298 10 of 23 www.geusbul let in.org (fig. 3). the age and lithological character of the succession in this area has been documented by cowie & adams (1957), frykman (1979), hambrey et al. (1989), smith & bjerreskov (1994) and smith & rasmussen (2008). no caves have been recorded to date elsewhere in the outcrop belt of the cambrian–ordovician. albert heim bjerge a sub-horizontal phreatic tube was observed in steeply dipping, floian (lower ordovician) subtidal limestones of the cape weber formation at the eastern end of albert heim bjerge, adjacent to wordie gletscher (fig. 3; 74.10764°n, 22.92733°w) and close to the palaeokarst documented in section 2.4. the cave is located c. 220 m above the base of the cape weber formation at 120 m a.s.l. a triangular, flat-based conduit (fig. 9) with a maximum width of 120 cm and height of 85 cm extends for 10 m to an ice choke, and there is no draft. the position of the cave, both stratigraphically and topographically, is similar to those that have a red-stained fill, and a similar timing of cave formation cannot be precluded in this instance. c. h. ostenfeld nunatak caves are located at the southern end of c. h. ostenfeld nunatak (fig. 3), overlooking grantagletscher, at an altitude of around 1000 m (74.28727°n, 22.89584°w). numerous phreatic conduits are erosionally truncated in a cliff of subtidal limestones of the heimbjerge formation (darriwilian–sandbian; middle to upper ordovician; smith & bjerreskov 1994) that dip gently eastwards (fig. 10). in addition to the large conduits described next, there are several c. 0.5 m diameter tubes in the vicinity. four main conduits are visible in the cliff (caves 1–4, fig. 10). cave 1 is a 12 m long, ascending phreatic tube (relative to the modern land surface) of around 1 m width (fig. 10a, b) that ends in a hoar-frost lined chamber that is c. 3 m in diameter and 1.5 m high (fig. 10c). cave 2 has a blocky breakdown breccia and, beyond a zone of frost shattering, a 0.5 m-wide, square-shaped conduit extends for 10 m to a choke. cave 3 has a 2 m high by 1 m wide conduit, wider at mid-height with a flat floor (fig. 10d), and again extends 10 m to an ice choke. cave 4 was not entered but is clearly seen above caves 2 and 3 in the cliff. the close spacing suggests that the individual caves may have been part of a single phreatic network prior to erosion. there is no trace of the cemented, coarse, red-stained fill that characterises the albert heim bjerge palaeokarst, and there is no evidence of any fill to these caves other than ice, uncemented silt and breakdown breccias that are also uncemented. fig. 9 a subhorizontal phreatic tube in the cape weber formation at the eastern end of albert heim bjerge adjacent to wordie gletscher (fig. 3). the triangular, flat-based conduit extends for 10 m to an ice choke. hammer (ringed) is 33 cm long. https://doi.org/10.34194/geusb.v49.8298 http://www.geusbulletin.org smith & moseley 2022: geus bulletin 49. 8298. https://doi.org/10.34194/geusb.v49.8298 11 of 23 www.geusbul let in.org together with these characteristics, the size and relative morphological complexity of these caves suggests that they are geologically younger than the palaeokarst described in albert heim bjerge. kronprins christian land the caves of kronprins christian land are located to the north and south of centrumsø (figs 2, 11). they are the best documented in greenland and have been visited by several expeditions. caves were first documented in the area by operation groundhog, as part of investigations into aircraft landing sites for the us air force (davies & krinsley 1960). twelve caves were reported in a large valley they named grottedal. the caves were next visited by a french caving expedition in 1983 (loubière 1987) and by geus geologists (including the first author) in 1994 and 1995 during the regional mapping programme. most recently, caves in grottedal and associated valleys have been extensively documented as part of the greenland caves project in 2015 and 2019 (moseley 2016; moseley et al. 2020), and for that reason, they are not described in detail here. caves in grottedal and its tributary valleys and one isolated example to the south of centrumsø (fig. 2) occur in the limestones of the odins fjord formation (llandovery, silurian) particularly in the vicinity of overlying carbonate mud mounds of the samuelsen høj formation (fig. 4; smith & rasmussen 2020). the odins fjord formation comprises grey to golden-brown weathering, highly fossiliferous lime mudstones and wackestones with coral-stromatoporoid biostromes. the mud mounds of the samuelsen høj formation are unbedded with no framework although rudstone flanking beds are commonly developed (smith & rasmussen 2020). the initiation of mud mound growth, which defines the formation boundary, commenced in the late llandovery (armstrong 1990) at around 436 ma. the grottedal caves are developed at several levels in the sub-horizontal limestones (moseley et al. 2020; smith & rasmussen 2020). u-shaped cave (fig.  11; ilusilik qaarusussuaq, gd4 of moseley 2016) was first described by davies & krinsley (1960) and is stratigraphically and topographically one of the lowest of the caves at around 100 m below the top of the odins fjord formation. several caves are distinctively developed at the odins fjord – samuelsen høj formation boundary such that the roofs of the caves are in reef lithofacies and the lower part of the conduit in bedded limestones of the lower unit. this is particularly well seen in cairn climb cave (inussuk innartooq qaarusussuaq, gd18) and crystal palace cave (aligoq illussaarsuaq qaarusussuaq, gd19), where large phreatic tubes are present at the boundary. other examples developed at the boundary are smaller but include multi-level cave (gd11) and triplet arch cave (gd21–23). detailed descriptions and surveys of these caves are available in moseley (2016) and moseley et al. (2020). b d c a 12 3 4 fig. 10 a complex of erosionally truncated phreatic tubes in subtidal limestones of the heimbjerge formation (darriwilian–sandbian, ordovician) at the southern end of c. h. ostenfeld nunatak, north-east greenland (fig. 3); looking north. a: general view of four of the entrances; the part of the cliff containing cave 1 is c. 10 m high. b: entrance to cave 1, which is 1.5 m high. c: conduit in cave 1, which extends for 12 m before terminating in a 3 m wide × 1.5 m high chamber with an ice choke; maximum conduit height is 2 m. d: entrance and conduit in cave 3 leading to ice choke. hammer on rucksack is 33 cm long. fig. 11 u-shaped cave (ilusilik qaarusussuaq, gd4 of moseley 2016) on the southern side of grottedal, kronprins christian land, north-east greenland (figs 1, 2). photo: robbie shone. https://doi.org/10.34194/geusb.v49.8298 http://www.geusbulletin.org smith & moseley 2022: geus bulletin 49. 8298. https://doi.org/10.34194/geusb.v49.8298 12 of 23 www.geusbul let in.org most of the conduit development in the kronprins christian land caves comprises sub-horizontal phreatic tubes. some of these are quite sizeable, and the largest conduit (in u-shaped cave) has a width of 8–13 m with a height of up to 10 m, variably filled with unlithified breakdown (moseley et al. 2020, fig. 2–2d). other phreatic features include anastomoses in the roof of kodak cave (gd8), together with phreatic scallops indicating flow inwards from the current entrance, towards the west (moseley 2016; moseley et al. 2020). few of the caves in the region, or in greenland more generally, have any vadose development. one exception is grotte des quatre in a thrust sheet of vertically bedded odins fjord formation to the south of centrumsø. this is the most geomorphologically complex cave in greenland with an upper level of phreatic morphology and a lower level that exhibits meandering vadose entrenchment (moseley et al. 2020, fig. 4–4c). a second example is swirly cave (sangujoraartoq qaarusussuaq) on the southern side of grottedal (moseley et al. 2020, fig. 10), which has a walking-size vadose canyon that extends downwards from the entrance to two large boulders, beyond which it narrows to a slot. a less well-developed example of vadose development is cove cave (eqik qaarusussuaq) in a canyon to the south of grottedal, where a 5 m deep vadose slot is developed in the floor of a linear phreatic conduit (moseley et al. 2020, fig. 13–13e). finally, kodak cave (gd8) contains three vadose wall notches (moseley 2016). a 12-cm-thick sample of flowstone was collected from cave gd8 in grottedal (moseley 2016; moseley et  al. 2021). u-th dating combined with an orbitally refined age model indicates that speleothem precipitation occurred between c. 588 and c. 549 ka (moseley et al. 2021), spanning the mis 15a–14 boundary and indicating a warmer, wetter climate in eastern north greenland at this time. speleogenesis must be older than the speleothem deposits, but several lines of evidence indicate that cave formation was much older than this. in particular, all of the explored cave conduits in kronprins christian land are present in vertical rock faces, consistent with the conduits being truncated by glacial erosion, which places older age constraints on the cave formation. central and western north greenland caves are abundant in the ordovician and, particularly, silurian limestones that extend westwards from j. p. koch fjord to hall land (figs 1, 8; supplementary file s1). fig. 12 phreatic tube in the basal kap jackson formation (sandbian, late ordovician) in south-eastern freuchen land, central north greenland, at the southern end of navarana fjord (figs 1, 8). the cave is one of several at this stratigraphic horizon at the locality and currently represents the most northerly documented cave. all penetrate for 5–7 m before terminating in a silt or ice choke. divisions on the survey pole are 20 cm. https://doi.org/10.34194/geusb.v49.8298 http://www.geusbulletin.org smith & moseley 2022: geus bulletin 49. 8298. https://doi.org/10.34194/geusb.v49.8298 13 of 23 www.geusbul let in.org they  are developed in topographically consistent positions, high in the vertical walls of the broadly n–s oriented fjords, indicating that cave formation predates incision by the outlet glaciers of the greenland ice sheet. most of the exposed conduits are of phreatic origin and are of very large scale, with conduit diameters ranging up to several tens of metres. details of the geographic and stratigraphic locations of the caves are given in table 1. freuchen land – 82.5°n caves were documented in wackestones and packstones at the base of the kap jackson formation (fig. 4; early sandbian, upper ordovician) in southern freuchen land as part of the regional mapping programme in 1984 (smith et al. 1989). the caves lie on the north side of a small glacier in south-eastern freuchen land that flows from the main ice sheet (82.31296°n, 41.83505°w), close to the southern end of navarana fjord (fl1 in fig. 8). a series of short phreatic tubes lie at the western end of (and 10 m above) the bench that marks the boundary between the cape webster and kap jackson formations. the tubes have a maximum length of 5–7 m, and the largest has a width of 2.3 m and a height of 1.6 m (fig. 12). one tube exhibits phreatic roof domes, and all are choked by either ochre-coloured fine sediment or ice. these are currently the northernmost documented karst caves globally. nordenskiöld fjord – 82°n a large entrance of around 30 m width was observed during helicopter reconnaissance in a 700 m high cliff at the southern end of nordenskiöld fjord (nf1 in fig. 8; 82.15836°n, 44.32837°w). the cave lies around 200 m below the cliff-top and is developed within a carbonate mound (m in fig. 13) immediately above the djævlekløften – petermann halvø formation boundary (ph in fig. 4). the mound has pale flanking beds with depositional dips markedly steeper than the inter-mound limestones. several mounds are visible at this horizon in nordenskiöld fjord, and dark grey to black, well-bedded lime mudstones are present between the mounds (higgins et al. 1991). the upper part of the entrance is a large phreatic conduit that extends into the cliff as a dark shadow above a ramp of talus. it is not clear to what extent the entrance has been modified by recent weathering, but a large notch is visible underneath the phreas, consistent with the development of a vadose notch, and is infilled by the talus ramp. wulff land – 81.8°n wul8 is a large cave entrance in wulff land at the north-western corner of apollo sø (figs 8, 14a, b; 81.84668°n, 48.19241°w; see also supplementary file s1). the entrance was first noted by davies & krinsley (1960) around 120 m below the top of the cliff and subsequently examined by helicopter during regional geological mapping in 1984. the cave entrance is approximately 30 m wide (fig. 14b) and a large talus ramp feeds a steep gulley and alluvial fan at the foot of the slope. the top of the talus ramp is level and the conduit can be seen to continue beyond. the cave occurs at a significant lithofacies boundary within the djævlekløften formation, where a cliff-forming unit principally composed of interbedded dark and light stromatoporoid biostromes with abundant crinoid debris (wg4 of sønderholm et al. 1987) overlies dark grey to black, thinly-bedded lime mudstones with occasional carbonate mounds, one of which occurs at the northern end of the hill (wg3 of sønderholm et al. 1987; see their fig. 5, which is the same cliff). a very large cave entrance, with associated minor entrances, wul6 (fig. 14c) was discovered as part of af kj cw ph dk m fig. 13 large, truncated cave conduit in the 700 m high wall of nordenskiöld fjord, central north greenland (figs 1, 8), with fine-grained fill weathering out of the entrance; the conduit is around 30 m wide and is located immediately above the djævlekløften – petermann halvø formation boundary, within a carbonate mound. af: aleqatsiaq fjord formation. cw: cape webster formation. dk: djævlekløften formation. kj: kap jackson formation. ph: petermann halvø formation. m: carbonate mounds of the djævlekløften formation. see fig. 4 for unit ages. https://doi.org/10.34194/geusb.v49.8298 http://www.geusbulletin.org smith & moseley 2022: geus bulletin 49. 8298. https://doi.org/10.34194/geusb.v49.8298 14 of 23 www.geusbul let in.org the regional mapping programme, 8 km to the southwest of wul8 on the west side of the large un-named valley that lies parallel to apollo sø and to the west of it (fig. 8; 81.78481°n, 48.47232°w), at c. 650 m a.s.l. the truncated phreatic tube is trapezoidal in transverse profile with a maximum width of around 20 m and a height of 10 m. inside, a ramp of debris extends upwards at around 30° but possible open cave could be seen at the top of the ramp. the cave is developed within sub-horizontal, burrow-mottled limestones of the aleqatsiaq fjord formation (smith et al. 1989), which is of mid-katian to llandovery (late ordovician–silurian) age (fig. 4). additional entrances were observed on oblique aerial photographs on the opposite side of the valley to wul6 and around 3 km south-west of wul8 (cave wul7, supplementary file s1, 81.82684°n, 48.32730°w). at least six entrances are developed along the base of a carbonate mound within the lower djævlekløften formation that dips gently north-west. the multiple entrances suggest that a phreatic network has been transected by glacial erosion. this cave development is unlikely to be related in speleogenetic terms to the wul6 conduits as it is developed at a much higher stratigraphic level, but wul7 does table 1 geographic and stratigraphic location of caves in central and western north greenland, together with image resources. caves latitude, longitude altitude plateau altitude stratigraphic unit photo number hall land (hl) hl1 81.32145, -57.32228 600 1000 petermann halvø formation nh-1985-b-032-008 hl2 81.30828, -57.31335 600 1000 djævlekløften formation nh-1985-c-011-004 nyeboe land (nl) nl1 81.17508, -56.78120 600 900 aleqatsiaq fjord formation nh-1984-b-025-019 nl2 81.26632, -57.05613 750 900 djævlekløften formation fig. 15e nl3 81.33289, -56.91269 600 900 djævlekløften formation nh-1984-c-008-022 nl4 81.53769, -57.13090 500 800 djævlekløften formation nh-1985-b-066-003 nl5 81.51205, -56.67436 600 800 dj®vlekløften formation nh-1985-b-065-027 nl6 81.51033, -55.86095 700 800 djævlekløften formation nh-1985-c-025-017 nl7 81.39275, -55.86233 500 900 petermann halvø formation nh-1985-c-025-012 nl8 81.36083, -55.97906 600 900 petermann halvø formation nh-1985-c-009-012 nl9 81.32756, -55.88430 700 900 petermann halvø formation nh-1985-c-009-010 nl10 81.30028, -55.82104 700 900 aleqatsiaq fjord formation nh-1985-c-009-009 nl11 81.63091, -54.83647 700 1100 djævlekløften formation fig. 15d warming land (wal) wal1 81.64952, -54.37494 600 1000 djævlekløften formation nh-1985-c-001-015 wal2 81.51402, -54.48893 500 900 petermann halvø formation nh-1985-b-002-001 wal3 81.50083, -53.45759 700 900 petermann halvø formation nh-1985-c-002-003 wal4 81.70141, -53.32850 600 1000 djævlekløften formation nh-1985-c-001-031 wal5 81.69487, -52.39740 700 1100 djævlekløften formation nh-1984-c-012-023 wal6 81.59510, -52.24085 700 900 aleqatsiaq fjord formation nh-1984-c-012-027 wal7 81.75259, -51.95520 500 800 djævlekløften formation nh-1985-c-008-019 wal8 81.73477, -51.83186 600 1000 petermann halvø formation fig. 15a wal9 81.64486, -51.54484 650 900 djævlekløften formation fig. 15b, c permin land (pl) pl1 81.76106, -51.52399 800 1000 djævlekløften formation nh-1984-c-012-010 wulff land (wul) wul1 81.81245, -50.37283 700 1000 djævlekløften formation nh-1985-c-004-003 wul2 81.71010, -50.21156 700 1000 djævlekløften + petermann halvø formations nh-1985-c-004-006 wul3 81.87180, -49.57921 700 1000 djævlekløften formation nh-1985-c-021-016 wul4 81.78847, -49.21392 450 1000 aleqatsiaq fjord formation nh-1985-b-010-027 wul5 81.84054, -48.61301 700 900 djævlekløften formation nh-1985-c-004-024 wul6 81.78480, -48.47232 650 900 aleqatsiaq fjord formation fig. 14c wul7 81.82684, -48.32730 700 900 djævlekløften formation nh-1985-c-004-025 wul8 81.84386, -48.18393 700 900 djævlekløften formation fig. 14a, b nordenskiold fjord (nf) nf1 82.15836, -44.32837 600 800 djævlekløften formation fig. 12 freuchen land (fl) fl1 82.30858, -41.85196 450 800 kap jackson formation fig. 13 further detail of these caves is provided in supplementary file s1, and for the location of land areas and caves, also see figs 1 and 8. ‘ plateau altitude‘ refers to the maximum elevation of the land surface (excluding ice) in the vicinity of the cave. image numbers with the prefix nh-198nrefer to oblique aerial photography of fjord walls obtained during the 1984–1985 regional mapping programme and available on the greenland mineral resources portal (http://maps.greenmin.gl/geusmap/?mapname=greenland_portal&lang=en). https://doi.org/10.34194/geusb.v49.8298 http://www.geusbulletin.org http://maps.greenmin.gl/geusmap/?mapname=greenland_portal&lang=en smith & moseley 2022: geus bulletin 49. 8298. https://doi.org/10.34194/geusb.v49.8298 15 of 23 www.geusbul let in.org have a very similar stratigraphic position to wul8. other caves observed  in wulff land during aerial reconnaissance (wul 1–5; table 1) are documented in supplementary file s1. warming land – 81.55–81.75°n several entrances are seen in the western wall of the fjord that separates permin land from southern warming land, in the innermost extension of hartz sund (fig.  8). cave wal8 (figs 8, 15a; 81.73477°n, 51.83186°w) is located c. 400 m below the top of an isolated 900 m hill. stratigraphically, it is developed within the djævlekløften formation at a prominent boundary between cliff-forming dark limestones and overlying pale carbonates. the conduit width is c. 20 m. multiple caves are observed in fjord and valley walls across warming land (wal1–8 in table 1) and permin land (pl1 in table 1) and are described in detail in supplementary file s1. a second group of cave entrances (wal9) is also visible in the fjord wall, 11 km to the south-southwest of the northerly cave (figs 8, 15b, c; 81.64486°n, a b c fig. 14 cave wul8, a large cave in the djævlekløften formation first noted by davies & krinsley (1960) at the north-western end of apollo sø, wulff land, western north greenland (figs 1, 8). a: cave from the east side of apollo sø. cliff is 800 m high from terrace to summit. b: close-up of cave entrance from helicopter; entrance approximately 30 m wide. c: the large cave entrance of wul6 photographed on a helicopter reconnaissance. located 8 km to the sw of wul8 at an altitude of c. 650 m. the truncated phreatic tube, developed within the aleqatsiaq fjord formation (fig. 4), is trapezoidal in transverse profile with a maximum width of around 20 m and a height of 10 m. https://doi.org/10.34194/geusb.v49.8298 http://www.geusbulletin.org smith & moseley 2022: geus bulletin 49. 8298. https://doi.org/10.34194/geusb.v49.8298 16 of 23 www.geusbul let in.org 51.54484°w) and 12 km north of the geological survey of greenland (ggu) base camp in 1984–1985. they occur 150 m below the  top of the 800 m cliff that marks the plateau  edge  and are again located in cliff-forming limestones of the djævlekløften formation. the conduit widths are less than 10 m, but the closely spaced cluster of multiple entrances again suggests erosional truncation of a phreatic network. cave wal6 is located within the interior of warming land (fig. 8, 81.59510°n, 52.24085°w; see also supplementary file s1) and is noteworthy for being one of the few caves in greenland with active water flow, with a stream emitting from the entrance at an  altitude of around 700 m (see supplementary file s1 for image). a stream appears to sink just back from the cliff edge and emerges from the entrance several tens of metres lower. a second entrance also has a suggestion of vertical development and is located a few tens of metres farther to the south. eastern nyeboe land – 81.6°n caves in the western wall of southern sankt george fjord, (nl11 in figs 8, 15d; table 1; 81.63091°n, 54.83647°w) were first recorded by niels henriksen in 1985 as part of the photo-reconnaissance flights to obtain oblique aerial photography of fjord and valley walls during the regional mapping programme (fig. 15d). the caves were subsequently photographed during a nasa operation icebridge mission (studinger 2011), and one of those is perhaps the largest cave passage identified in greenland to date. an entrance of c. 30 m width is located within sub-horizontal limestones of the djævlekløften formation. it has a talus fan forming the floor, vertical walls and a gently arched roof, creating an overall a b d e c fig. 15 caves exposed in fjord walls incised by outlet glaciers from the greenland ice sheet in warming land and nyeboe land, north greenland. a: cave entrance (wal8) at the base of the petermann halvø formation (llandovery, silurian) where it overlies the aleqatsiaq fjord formation in eastern warming land (figs 1, 8). cliff is 900 m high, and the cave entrance is c. 20 m wide. b: multiple cave entrances (arrowed; wal9) within djævlekløften formation, 11 km to ssw of fig. 14a, where a phreatic network has been transected by glacial erosion. cliff is 800 m high, and entrances are less than 10 m in diameter. c: enlargement of multiple entrances seen arrowed in b. d: multiple cave entrances (nl11) developed in the djævlekløften formation exposed in a cliff in south-eastern nyeboe land, western north greenland (fig. 1). cliff is 1100 m high from fjord to summit, and the larger entrance is 30 m wide. e: cave nl2 in south-western nyeboe land in cliff bordering the glacier at the head of newman bugt. the cave is located in the upper aleqatsiaq fjord formation and the cliff is 650 m high from glacier to plateau edge. the entrance is around 15 m wide, and the wash of fine sediment indicates that there is intermittent water flow through the cave. photos: niels henriksen (a, nh-1985-c-008-018; b, c, nh-1985-c-008-0115; d, nh-1985-c-011-033; e, nh-1984-c-008-034). https://doi.org/10.34194/geusb.v49.8298 http://www.geusbulletin.org smith & moseley 2022: geus bulletin 49. 8298. https://doi.org/10.34194/geusb.v49.8298 17 of 23 www.geusbul let in.org sub-quadrate conduit cross-section. a second smaller entrance is present 130 m to the north at a similar stratigraphic horizon and may represent the same conduit or system. the caves are located c. 360 m below the top of the cliff that rises to 1000 m from sea level. the caves lie directly beneath carbonate mounds of the hauge bjerge formation, but unlike other examples in north greenland they are at a considerably lower level, around 130 m below the reef mounds, and there is unlikely to be a genetic association. south-western nyeboe land and hall land multiple cave entrances are present in the west-facing walls that bound the glacier feeding the head of newman bugt (nl2 in figs 8, 15e; table 1; 81.26632°n, 57.05613°w). the most prominent of these (fig. 15e) has a 15 m wide entrance located 100 m below the edge of the plateau surface at 900 m. it occurs within the aleqatsiaq fjord formation, just below the boundary with the overlying petermann halvø formation (fig. 4) and is one of relatively few caves developed within that unit, perhaps because the rubbly weathering nature does not lend itself to stable conduit formation. the cave, and others along this stretch of cliffs, has a conspicuous wash of sediment descending from the entrance, suggesting that water flows out from the cave entrance at least intermittently. caves are abundantly developed in the southern, platform areas of nyeboe land and hall land but to date have only been observed during aerial reconnaissance flights. a summary of caves hl1–2 and nl1–nl11, identified from oblique aerial photography, is available in supplementary file s1. age constraints on cave formation in north and north-east greenland determination of the timing of speleogenesis for the large, open caves exposed in the fjords and cliffs of north-east and north greenland is dependent on being able to distinguish younger karst features from older palaeokarst. pre-cenozoic palaeokarst is extensive across north america and is particularly well developed at the tops of the sauk megasequence (middle ordovician) and the kaskaskia megasequence (late mississippian) of sloss (1963), driven by large-scale changes of relative sea level in carbonate-dominated successions (palmer & palmer 1989, 2011). extensive palaeokarst has also been documented in the middle devonian and to a lesser degree late devonian of canada (ford 1989). these occurrences of endokarst in palaeozoic rocks across north america are preserved in the geological record by the cavities being infilled with lithified younger sediments that prevent cave collapse. if unfilled cavities undergo deep burial then collapse and mechanical compaction will occur, with a zone of suprastratal deformation extending upwards from the collapse; by the time burial has proceeded to 2–3 km, few cavities are detected by drill-bit drops during drilling (loucks 2007). the large cave conduits documented in north and north-east greenland lack a lithified fill (in distinct contrast to the palaeozoic palaeokarst), despite being environments with an abundance of calcium carbonate to form cements for any cave sediments. despite their large size and open character, they also lack large-scale collapse and suprastratal deformation; where collapse occurs, it is localised and probably related to frost-shattering. together these characteristics indicate that the caves are geologically young, and it is highly unlikely that the caves have been buried to any significant depth or been through multiple burial–uplift cycles. these geologically young caves, in contrast to the older palaeokarst, consistently occur within the upper few hundred metres beneath the present-day topographic surface, where they are exposed in the fjord walls and other cliffs. speleogenesis must therefore post-date formation of a terrestrial surface following uplift but predate the earliest fjord-forming glaciation by outlet glaciers from the greenland ice sheet that overprinted the landscape. onset of glaciation in north and north-east greenland the earliest evidence for cenozoic glaciation in greenland is the presence of ice-rafted debris off the eastern coast of greenland at 75°n derived, at least intermittently, from glaciation during the late eocene and early oligocene (30–38 ma) that probably corresponded to alpine-style glaciers (eldrett et al. 2007, 2009). during the cenozoic, northern hemisphere glaciation is unlikely to have been possible prior to the late eocene (deconto et al. 2008; bierman et al. 2016). episodic glaciation off the coast of south-eastern greenland through the miocene and pliocene is evidenced by glacial dropstones and ice-rafted debris from 7 ma (late miocene; larsen et al. 1994) and quartz sand grains with surface textures characteristic of glacial erosion (helland & holmes 1997; see bierman et al. 2016, fig. 5 for summary). modelling of the uplift history and glaciation by solgaard et al. (2013) showed an intimate connection between uplift and glaciation by combining geological observations and climate modelling. solgaard et al. (2013) concluded that the low-relief miocene surface prior to uplift at 10 ma and 5 ma (see below) had no ice. a cooling in surface temperature combined with an increase in orographic precipitation after 10 ma led to ice build-up although evidence suggests that this was confined to localised ice caps in northernmost and south-eastern greenland. a second period of uplift https://doi.org/10.34194/geusb.v49.8298 http://www.geusbulletin.org smith & moseley 2022: geus bulletin 49. 8298. https://doi.org/10.34194/geusb.v49.8298 18 of 23 www.geusbul let in.org at 5 ma enabled ice sheet initiation and relocated the main area of ice sheet nucleation from northernmost greenland to the south-east, south and south-west. during colder-than-present climatic conditions, ice could flow into the interior of greenland and form an inland ice sheet although this process was inhibited by a föhn effect (solgaard et al. 2013, fig. 13). nevertheless, due to climatic deterioration, an expansive and persistent greenland ice sheet had developed across the interior by c. 2.7 ma (flesche kleiven et al. 2002; bierman et al. 2016). the first physical evidence of glaciation preserved in north greenland is a till that is overlain by nonglacial, marine and deltaic sediments of the kap københavn formation (funder et al. 2001), which are in turn overlain by another till. the difficulty of dating the kap københavn formation has resulted historically in a narrow spread of ages around c. 2.4 ma (funder et al. 2001), but bennike et  al. (2010) undertook a detailed re-appraisal of the stratigraphic correlation of pliocene– pleistocene units in north and north-east greenland. the kap københavn formation was separated into two disconformable members, of which the younger one, member b, was thought to have been deposited just prior to the base of the olduvai subchron (1.95 ma). the older member a, beneath the disconformity, was interpreted as being of terminal pliocene age (c. 2.6 ma). funder et al. (2001) considered the underlying glacial till to be in conformable contact with the overlying nonglacial sediments and the oldest dateable field evidence for the initiation of glaciation in north greenland, which could have begun to incise the plateau and transect the caves is therefore c. 2.7 ma. independent evidence that the climate remained intermittently warm into the earliest pleistocene, at temperatures sufficient to support deciduous vegetation in north greenland, comes from sporadic occurrences of sub-fossil wood across north greenland, dated at c. 3 ma (bennike 1998, 2000); the preservation of larix and thuja in member b of the kap københavn formation (2 ma) (funder et al. 2001); and a variety of deciduous land plants, including larix and betula, in the store koldewey formation of north-east greenland (1.95–1.78 ma; bennike et al. 2010). given the paucity of sedimentary evidence for the onset of fjord-forming glaciation in north greenland, pedersen et al. (2019) took a geophysical approach by examining the flexural isostatic response to erosional unloading around independence fjord (fig. 1), concluding that the fjord systems in this region must have formed by glacial erosion at average rates of c. 0.5–1 mm yr-1 since around 2.5 ma. it may therefore be concluded that pre-existing phreatic cave systems were truncated by glacial erosion by the latest pliocene or earliest pleistocene and that there is no evidence for active conduit formation or enlargement after that date although speleothem deposition has occurred within open cavities since then (moseley et al. 2021). age constraints of cave formation on albert heim bjerge and c. h. ostenfeld nunatak the caves developed in the cape weber formation on albert heim bjerge are short (10 m), at low altitude (100–200 m a.s.l.), and are morphologically simple phreatic tubes similar to the palaeokarst caves infilled with devonian molasse. although they could be geologically young features, it cannot be precluded that they are contemporaneous with those caves and erosional pockets nearby that preserve a devonian fill. in contrast, the caves developed on c. h. ostenfeld nunatak are morphologically more complex, as evidenced by the multiple entrances to what might have been part of a single phreatic endokarst complex and also by the more varied conduit geometries. the presence of open conduits without lithified sedimentary infill or significant collapse strongly suggests that the c. h. ostenfeld nunatak caves are geologically young morphological features, in contrast to those filled with devonian molasse on albert heim bjerge. the caves on c. h. ostenfeld nunatak are located at the southern end of the nunatak at an altitude of 1000 m a.s.l. and are coincident with an observed erosional remnant of a neogene peneplanation surface, termed the lower planation surface (lps), that forms the highest ground on the nunatak (bonow & japsen 2021, figs 7, 9, supplementary file s2). given that the caves predate late pliocene glaciation, they may have been associated with the uplift events that formed successive planation surfaces along the eastern coast of greenland from at least 68°n to 78°n (bonow & japsen 2021). these authors hypothesised that the region was subaerially eroded down to a surface of low topographic relief at a marine base level around the eocene–oligocene boundary, which they termed the upper planation surface (ups). the ups was subsequently uplifted at c. 10 ma (late miocene) by around 1 km, leading to incision towards base level and the development of the lps. further uplift at c. 5 ma (early pliocene) then elevated the lps by a further 1 km above the base level. given that the caves on c. h. ostenfeld nunatak are geologically young, it is possible that they are linked genetically with these neogene uplift events. age constraints of cave formation in north greenland the area of the laurentian shelf in north greenland in which caves have been recorded, from kronprins https://doi.org/10.34194/geusb.v49.8298 http://www.geusbulletin.org smith & moseley 2022: geus bulletin 49. 8298. https://doi.org/10.34194/geusb.v49.8298 19 of 23 www.geusbul let in.org christian land in the east (section 3.2) to hall land in the west (section 3.3; fig. 1), is characterised by a plateau at 800–1000 m, into which fjords with vertical walls are deeply incised by outlet glaciers from the greenland ice sheet. this plateau surface is highest in the southern parts of nyeboe land and warming land, reducing in height eastwards into wulff land and the area around the head of victoria fjord (fig. 8). it is high again in central peary land, at around 1000 m, but mainly at lower altitudes in kronprins christian land. there are two areas of conspicuously higher elevation, one to the north of peary land in the mountains of johannes v. jensen land (up to 1850 m a.s.l.) and the second in the prinsesse caroline-mathilde alper of central kronprins christian land (up to 1742 m a.s.l.; japsen et al. 2021, fig. 19). the type of landscape analysis that has elucidated the spatial distribution of the ups and lps on the eastern and western coasts of greenland has not yet been extended to north greenland, but the region does have the characteristics of an elevated passive continental margin sensu green et al. (2013, 2018). apatite fission track analysis (afta), together with some vitrinite reflectance data, is, however, available to constrain the cenozoic uplift history around independence fjord and the margins of the wandel sea basin (japsen et al. 2021), between the two main cave-bearing regions of kronprins christian land and western north greenland (fig. 1). the digital elevation model for the independence fjord region (japsen et al. 2021, fig. 19) shows that there is a plateau surface on either side of independence fjord at around 1000 m a.s.l., and it is this surface that is continuous with the plateau developed at a similar altitude across western north greenland (fig. 8). in relation to cenozoic uplift, the afta data identified a short, localised phase of uplift in the mid-paleocene (60 ma) associated with inversion and exhumation along major fault zones, and a more regional, but again short lived, cooling/uplift event at the end of the eocene that led to the almost complete removal of an extensive, kilometer-thick sedimentary cover deposited during eocene subsidence (piasecki et al. 2018; japsen et al. 2021). the afta data also identified an interval of mid-late miocene cooling and uplift in the interval from 15–5 ma that produced the modern plateau topography (japsen et al. 2021), followed by incision. this event is not associated with any known sedimentary record, but japsen et al. (2021) did note that this uplift event does correlate broadly with the events in north-east greenland that generated the ups and lps. on the balance of evidence, japsen et al. (2021) considered it most likely that the north greenland plateau surface correlates temporally with the initial miocene (10 ma) uplift of the ups farther to the south, rather than the 5 ma pliocene event that uplifted the lps and further elevated the ups. the lack of either fill or collapse in the caves developed in north greenland indicates that they were not part of the eocene subsidence and basin fill associated with the deposition of the thyra ø formation (piasecki et al. 2018). the oldest that the caves could be is therefore latest eocene, associated with the latest stages of exhumation that removed the eocene sedimentary cover identified by japsen et al. (2021). more likely, given their open character, lack of sedimentary fill, and absence of collapse, they are miocene and broadly associated with the uplift event that generated the plateau. the formation of very large-scale phreatic cave systems must have coincided with a period of relatively wet climate in northernmost greenland, and solgaard et al. (2013) did note that uplift of the low-relief surface during the miocene led to an orographic increase in precipitation after 10 ma. speleogenesis persisted no later than the latest pliocene or earliest pleistocene when the onset of major glaciation led to the formation of the fjords by valley outlet glaciers emanating from the greenland ice sheet (pedersen et al. 2019) and the erosional transection of the large-scale phreatic systems. conclusions spatially extensive cambrian–silurian carbonate rocks were deposited on the north-east corner of the laurentian palaeocontinent as part of the great american carbonate bank (derby et al. 2012), which is today represented by the north and east coasts of greenland, but which also constituted an original continental promontory during the early palaeozoic. these carbonates have been prone to the development of karst during times of relative sea-level lowstand. palaeokarst surfaces are frequently developed at the sequence boundaries with particularly well-developed examples in the portfjeld formation (late ediacaran – cambrian), beneath the buen and kap holbæk formations (cambrian series 2) of north greenland and beneath the post-orogenic devonian cover in southern north-east greenland. for the most part, this involves the development of morphologically simple karst with grikes (kluftkarren) infilled with sediment, which is often mature quartz sand that has been subject to aeolian reworking. however, in kronprins christian land (sæfaxi elv and hjørnegletscher, figs 1, 2) the palaeokarst includes infilled cave systems that include a variety of phreatic conduits of probably early cambrian age, together with penecontemporaneous surface channels that postdate the caves and truncate them. infilled palaeo-cave systems are also present in the lower ordovician cape weber formation on albert heim bjerge (fig. 6), and in this case, the sediment fills are of devonian age. caves are also widely developed, but for the most part have not previously been documented, in the https://doi.org/10.34194/geusb.v49.8298 http://www.geusbulletin.org smith & moseley 2022: geus bulletin 49. 8298. https://doi.org/10.34194/geusb.v49.8298 20 of 23 www.geusbul let in.org ordovician and silurian limestones of north-east and east greenland. in north greenland, where most of the caves occur, the shelf successions of this age are sub horizontal and have remained so since deposition. notably, these caves are similar to each other but contrast markedly with the older palaeokarst. for example, they lack a cemented sedimentary fill and the largest examples are of much larger scale, with conduits measuring tens of metres rather than a few metres or smaller. the only sedimentary fill within them is uncemented mud and silt and localised, uncemented breakdown breccias; many are plugged with ice. they also have a characteristic geomorphological position, located in the steep fjord walls and other glacially eroded cliffs, within a few 100 metres vertically beneath the plateau surface that is developed at elevations of 800–1000 m (fig. 8). given these similarities, it is probable that they have a similar speleogenetic history and formed in a single phase of cave development. in central and western north greenland, much of this phreas is very large, with individual conduits commonly 20–30 m wide, but where exploration has been more comprehensive, for example in kronprins christian land (moseley 2016; moseley et al. 2020), abundant smaller caves, still predominantly phreatic, have also been discovered. all caves explored to date end in loose sediment or ice chokes after a few tens of metres. the longest explored cave to date is cove cave in grottedal, kronprins christian land, with a surveyed length of 103 m (moseley et al. 2020). although short, the caves are significant in containing speleothems, which offer considerable potential for documenting high-latitude pleistocene palaeoclimate (moseley et al. 2021). in north greenland, given that the caves are most commonly exposed in fjord walls and other cliffs, they must predate incision by the outlet glaciers from the greenland ice sheet that created the fjords and truncated the functional cave systems. the earliest physical evidence for glaciation in north greenland is a till that conformably underlies marine sediments of the kap københavn formation (funder et al. 2001), which has been dated as the latest pliocene (bennike et al. 2010). modelling of early ice sheet growth in north and northeast greenland is highly responsive to changes in input parameters, and predictions of ice cover vary accordingly (solgaard et al. 2013), but studies of the flexural isostatic response to erosional unloading indicate that fjord formation in north greenland is likely to have commenced at around 2.5 ma (pedersen et al. 2019). the available evidence from the geological and geomorphological context thus indicates that the caves of north and north-east greenland ceased to be active by the latest pliocene or earliest pleistocene. the only activity within the caves after that date was the deposition of speleothems during warm, moist intervals (moseley et al. 2021), together with some local breakdown and the accumulation of fine, probably aeolian, sediment. the maximum age of formation is less well constrained in both north and north-east greenland. on c. h. ostenfeld nunatak, multiple open conduits suggest that a phreatic network has been truncated and passages are up to 3 m in width. successive uplift episodes at 10 ma and 5 ma created distinctive paired planation surfaces at c. 1 km and 2 km a.s.l. termed the lps and ups, respectively (japsen et al. 2014; bonow & japsen 2021). the cave systems observed on c. h. ostenfeld nunatak, north-east greenland (figs 3, 9), lie directly beneath a glacially eroded remnant of the lps (bonow & japsen 2021, figs 7, 9). the topographical context suggests that the cave formation was possibly associated with the miocene uplift. the earliest date for cave formation in north greenland is also poorly constrained, but the lack of cemented fill and of cave collapse shows that the caves have not been buried deeply within sedimentary basins and nor have they been through multiple cycles of burial–uplift. the clear association with the plateau surface provides a further line of evidence. in north greenland, there is only a single planation surface, which japsen et al. (2021) ascribed to a miocene uplift event at 10 ma (temporally equivalent to the ups on the east and west greenland coasts). this event creates the distinctive plateau surface at 800–1000 m from hall land in the west to kronprins christian land in the east (figs 1, 8). the combination of evidence suggests that the initiation of cave formation in north greenland is unlikely to be pre-miocene, and that speleogenesis was probably associated with miocene uplift. the association of phreatic systems with large conduits and their development beneath a low gradient coastal peneplain may suggest that the cave systems of the nullarbor plain in southern australia (webb & james 2006; woodhead et al. 2019) offer a partial analogue despite the contrast in climate systems. there, the larger caves that have long flow paths of 100 km or more developed up to 150 m below the low-relief land surface, and the conduit size is directly proportional to the length of the flow path, with deep conduit formation favoured for flow paths with lengths >3 km (webb & james 2006). the nullarbor systems are also geologically old, with most of the conduit development having occurred in the oligocene and only small amounts of speleogenesis since then, during the pliocene wet period in australia. the caves of north and north-east greenland, dissected by later erosion, offer a tantalising glimpse of a large-scale phreatic drainage system that developed beneath the water table and below a coastal peneplain during neogene time. https://doi.org/10.34194/geusb.v49.8298 http://www.geusbulletin.org smith & moseley 2022: geus bulletin 49. 8298. https://doi.org/10.34194/geusb.v49.8298 21 of 23 www.geusbul let in.org the caves of greenland preserve an important part of the geological history and landscape development of north greenland that is otherwise absent from the physical geological record. in north greenland, the youngest pre-glacial sediments that are preserved are the clastic sediments and coals of the thyra ø formation, which is of paleocene – middle eocene age (c. 56–41 ma; lyck & stemmerik 2000; piasecki et al. 2018). the caves provide physical evidence of a cenozoic, probably miocene, geological history that is otherwise unrecorded, and the study of the loose sediment fills may further elucidate this unrepresented and elusive period in the sedimentary record of north greenland. acknowledgements the fieldwork for this paper was mainly undertaken as part of the ggu/geus regional mapping programme in central north greenland (1984), kronprins christian land (1994, 1995) and north-east greenland (1988). other cave locations have been determined using the oblique aerial photography available in the greenland portal (http:// maps.greenmin.gl/geusmap/?mapname=greenland_portal&amp). the polar geospatial center, university of minnesota, is thanked for access to arctic dem data, and lena friedrich is thanked for assistance with gis. work undertaken as part of the greenland caves project (2015, 2019) was funded by the austrian science fund (project no. y 1162-n37 to moseley). the greenland government is thanked for permission to undertake this fieldwork (knno expedition permit c-19-32; scientific survey licence vu-00150; greenland national museum and archives 2019/01). we thank the reviewers for detailed and perceptive reviews that significantly improved the manuscript. additional information funding statement this work was funded by geus (mps fieldwork 1984–1998), carlsbergfondet and the agouron institute (mps fieldwork 2009, 2011) and the austrian science fund (y 1162-n37 to gem). competing interests the authors declare no competing interests. author contributions mps: conceptualisation, investigation (equal), writing – original draft (lead), writing – review and editing (equal). gem: investigation (equal), writing – original draft (supporting), writing – review and editing (equal). additional files a supplementary file (s1) is available in the geus bulletin data repository that provides a full record of caves identified in western north greenland, with additional images of these caves: https://doi. org/10.22008/fk2/gqgelo references armstrong, h.a., 1990: conodonts from the upper ordovician–lower silurian carbonate platform of north greenland. bulletin grønlands geologiske undersøgelse 159, 151 pp. https://doi.org/10.34194/bullggu.v159.6709 bengaard, h.j. & henriksen, n. 1986: geological map of greenland, 1:500 000, sheet 8, peary land. copenhagen: geological survey of greenland. https://doi.org/10.22008/fk2/q7hidy bennike, o. 1998: late cenozoic wood from washington land. geological survey bulletin 180, 155–158. https://doi.org/10.34194/ggub. v180.5100 bennike, o. 2000: notes on the late cenozoic history of the washington land area, western north greenland. geological survey bulletin 186, 29–34. https://doi.org/10.34194/ggub.v186.5212 bennike, o., knudsen, k.l., abrahamsen, n., böcher, j., cremer, h. & wagner, b. 2010: early pleistocene sediments on store koldewey, northeast greenland. boreas 39, 603–619. https://doi. org/10.1111/j.1502-3885.2010.00147.x bierman, p.r., shakun, j.d., corbett, l.b., zimmerman, s.r. & rood, d.h. 2016: a persistent and dynamic east greenland ice sheet over the past 7.5 million years. nature 540, 256–260. https://doi.org/10.1038/ nature20147 bonow, j.m. & japsen, p. 2021: peneplains and tectonics in north-east greenland after opening of the north-east atlantic. geus bulletin 45(1), 5297. https://doi.org/10.34194/geusb.v45.5297 cohen, k.m., finney, s.c., gibbard, p.l. & fan, j.-x. 2013: the ics international chronostratigraphic chart (updated july 2021). episodes 36. 199–204. http://www.stratigraphy.org/icschart/chronostratchart2021-07.pdf cowie, j.w. & adams, p.j. 1957: the geology of the cambro-ordovician rocks of east greenland. meddelelser om grønland 153(1), 193 pp. davies, w.e. & krinsley, d.b. 1960: caves in northern greenland. national speleological society bulletin 22, 114–116. dawes, p.r., glendal, e.w. & holst, j. 2016: a glossary for geus publications: spelling and usage of troublesome words and names made easy. danmarks og grønlands geologiske undersøgelse rapport 2016/3, 74 pp. deconto, r., pollard, d., wilson, p., pälike, h., lear, c.h. & pagani, m. 2008: thresholds for cenozoic bipolar glaciation. nature 455, 652– 656. https://doi.org/10.1038/nature07337 derby, j.r., raine, r.j., smith, m.p. & runkel, a.c. 2012: paleogeography of the great american carbonate bank of laurentia in the earliest ordovician (early tremadocian): the stonehenge transgression. in: derby, j.r. et al. 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https://doi.org/10.34194/rapggu.v143.8055 https://doi.org/10.1144/gsjgs.156.1.0113 https://doi.org/10.1144/gsjgs.156.1.0113 https://doi.org/10.1016/j.palaeo.2013.09.019 https://doi.org/10.34194/rapggu.v133.7973 https://doi.org/10.34194/rapggu.v133.7973 https://michaelstudinger.smugmug.com/greenland-2011/i-jrmsrzm https://michaelstudinger.smugmug.com/greenland-2011/i-jrmsrzm https://doi.org/10.1017/s001675680003572x https://doi.org/10.1130/0016-7606(1972)83 https://doi.org/10.1130/0016-7606(1972)83 https://doi.org/10.1130/dnag-gna-e https://doi.org/10.1130/dnag-gna-e https://doi.org/10.1130/2006.2404(07 https://doi.org/10.1038/s42003-020-01381-7 https://doi.org/10.1038/s42003-020-01381-7 https://doi.org/10.1038/s41598-018-37097-2 https://doi.org/10.1038/s41598-018-37097-2 the karst and palaeokarst of north and north-east greenland physical records of cryptic geological intervals abstract introduction palaeokarst in greenland intra-ediacaran palaeokarst sub-cambrian series 2 palaeokarst peary land 82.2-82.8°n kronprins christian land 80-81°n hypogene karst in kronprins christian land sub-devonian palaeokarst wordie gletscher, east greenland, 75°n caves of north and north-east greenland wordie gletscher, east greenland 74°n albert heim bjerge c. h. ostenfeld nunatak kronprins christian land central and western north greenland freuchen land 82.5°n nordenskiöld fjord 82°n wulff land 81.8°n warming land 81.55-81.75°n eastern nyeboe land 81.6°n south-western nyeboe land and hall land age constraints on cave formation in north and north-east greenland onset of glaciation in north and north-east greenland age constraints of cave formation on albert heim bjerge and c. h. ostenfeld nunatak age constraints of cave formation in north greenland conclusions acknowledgements additional information funding statement competing interests author contributions additional files references figures fig. 1 map of north greenland showing the outcrop of sandbian (upper ordovician) to wenlock (silurian) shelf limestones. palaeokarst localities are indicated by a red dot and cave localities by a black dot; dots may represent more than one instance where localities are closely spaced. for full details of all known caves, see supplementary information. fl: freuchen land. hg: hjørnegletscher. nb: newman bugt. naf: navarana fjord. nof: nordenskiöld fjord. øm: øvre midsommersø; se: sæfaxi elv. sp: sirius passet. sg: steensby gletscher. linework based on the 1:500 000 ggu and geus geological maps (bengaard & henriksen 1986; henriksen 1989; jepsen 2000). fig. 2 geological map of kronprins christian land, north-east greenland, showing the distribution of sandbian (upper ordovician) to wenlock (silurian) shelf limestones in which caves are developed (black dots). a regionally extensive palaeokarst horizon is preserved fig. 3 geological map of albert heim bjerge and c. h. ostenfeld nunatak showing cave (black dots) and palaeokarst (red dot) localities fig. 4 correlation chart of cambrian–silurian shelf units in central and western north greenland (nyeboe land – freuchen land), eastern north greenland (kronprins christian land) and north-east greenland (c. h. ostenfeld nunatak) showing the stratigraphic distribution of caves and palaeokarst. stratigraphic units are of formation rank unless otherwise stated. the left-hand columns follow the international union of geological science/international commission on stratigraphy (iugs/ics) standard (cohen et al. 2013); boundaries of megasequences and supersequences are from palmer (1981), golonka & kiessling (2002) and smith & rasmussen fig. 5 sub-cambrian palaeokarst developed within the upper 2 m of the portfjeld formation (ediacaran or lowermost cambrian) at the sirius passet lagerstätte locality, north-western peary land (fig. 1; harper et al. 2019) and infilled with clastic sediments of the buen formation (cambrian, series 2). a: steeply inclined fissure infilled by millet-seed quartz arenite, margins arrowed, with a fin of dolostone dividing the fissure into two close to the disconformity surface (skyline of block). b: inclined fissure (arrow) and sub-horizontal fig. 6 sub-cambrian palaeokarst developed within tonian (neoproterozoic) carbonate rocks of the fyns sø formation (fs) of kronprins fig. 7 examples of quadrate phreatic conduits (a, b) developed within subtidal limestones of the cape weber formation (tremadocian–floian, lower ordovician) infilled with devonian conglomerates of the vilddal group (middle devonian) that mainly comprise blocks of red-stained ordovician carbonates in a clastic matrix; the eastern end of albert heim bjerge adjacent to wordie gletscher (fig. 3). divisions on the survey pole are 20 cm, and lens cap is 7 cm. fig. 8 digital elevation model (dem) of central and western north greenland showing the caves described and figured in the text (data from arctic dem, porter et al. 2018). the caves are developed immediately beneath the plateau surface, at an altitude of 800–1000 m, that is seen in the southern parts of the area, from freuchen land westwards to hall land. the dark blue peaks are ice caps sitting on the plateau surface. documentation and illustration of other caves discovered during aerial reconnaissance are provided in supplementary fig. 9 a subhorizontal phreatic tube in the cape weber formation at the eastern end of albert heim bjerge adjacent to wordie gletscher (fig. 3). the triangular, flat-based conduit extends for 10 m to an ice choke. hammer (ringed) is 33 cm long. fig. 10 a complex of erosionally truncated phreatic tubes in subtidal limestones of the heimbjerge formation (darriwilian–sandbian, fig. 11 u-shaped cave (ilusilik qaarusussuaq, gd4 of moseley 2016) on the southern side of grottedal, kronprins christian land, north-east greenland (figs 1, 2). photo: robbie shone. fig. 12 phreatic tube in the basal kap jackson formation (sandbian, late ordovician) in south-eastern freuchen land, central north greenland, at the southern end of navarana fjord (figs 1, 8). the cave is one of several at this stratigraphic horizon at the locality and currently represents the most northerly documented cave. all penetrate for 5–7 m before terminating in a silt or ice choke. divisions on the survey pole are 20 cm. fig. 13 large, truncated cave conduit in the 700 m high wall of nordenskiöld fjord, central north greenland (figs 1, 8), with fine-grained fill weathering out of the entrance; the conduit is around 30 m wide and is located immediately above the djævlekløften fig. 14 cave wul8, a large cave in the djævlekløften formation first noted by davies & krinsley (1960) at the north-western end of apollo sø, wulff land, western north greenland (figs 1, 8). a: cave from the east side of apollo sø. cliff is 800 m high from terrace to summit. b: close-up of cave entrance from helicopter; entrance approximately 30 m wide. c: the large cave entrance of wul6 photographed on a helicopter reconnaissance. located 8 km to the sw of wul8 at an altitude of c. 650 m. the truncated phreatic tube, developed within the aleqatsiaq fjord formation (fig. 4), is trapezoidal in transverse profile with a maximum width of around 20 m and a height of 10 m. fig. 15 caves exposed in fjord walls incised by outlet glaciers from the greenland ice sheet in warming land and nyeboe land, north greenland. a: cave entrance (wal8) at the base of the petermann halvø formation (llandovery, silurian) where it overlies the aleqatsiaq fjord formation in eastern warming land (figs 1, 8). cliff is 900 m high, and the cave entrance is c. 20 m wide. b: multiple cave entrances (arrowed; wal9) within djævlekløften formation, 11 km to ssw of fig. 14a, where a phreatic network has been transected by glacial erosion. cliff is 800 m high, and entrances are less than 10 m in diameter. c: enlargement of multiple entrances seen arrowed in b. d: multiple cave entrances (nl11) developed in the djævlekløften formation exposed in a cliff in south-eastern nyeboe land, western north greenland (fig. 1). cliff is 1100 m high from fjord to summit, and the larger entrance is 30 m wide. e: cave nl2 in south-western nyeboe land in cliff bordering the glacier at the head of newman bugt. the cave is located in the upper aleqatsiaq fjord formation and the cliff is 650 m high from glacier to plateau edge. the entrance is around 15 m wide, and the wash of fine sediment indicates that there is intermittent water flow through the cave. photos: niels henriksen (a, nh-1985-c-008-018; b, c, nh-1985-c-008-0115; d, nh-1985-c-011-033; e, nh-1984-c-008-034). table table 1 geographic and stratigraphic location of caves in central and western north greenland, together with image resources. geological survey of denmark and greenland bulletin 11, 125-144 125 presentation and interpretation of structural data from the nagssugtoqidian orogen using a gis platform: general trends and features jeroen a.m. van gool and sandra piazolo in this contribution we present data collected by more than 50 international geologists involved in geological mapping and research projects in the nagssugtoqidian orogen of west greenland, organised by the geological survey of denmark and greenland and the danish lithosphere centre. using a geographical information system (gis) as a framework for visualisation and analysis of structural and lithological data, it is now possible to give a unique overview of thousands of data points, employed here within a study area of approximately 160 × 180 km in the central and northern nagssugtoqidian orogen. the gis methodology allows comparison, integration and analysis of datasets in terms of subject, space, and scale. this is extremely helpful in the recognition of geological patterns, such as terrain or domain boundaries and map-scale structures. analysis of the available structural data shows clear differences in deformation patterns between the core and the northern segment of the nagssugtoqidian orogen. one of the most prominent features is the ene-striking nordre strømfjord shear zone, which transects the orogen from the coast to the inland ice. the data also clearly document a change from predominantly steeply dipping, ene–wsw-trending fabrics and large, elongate structural domains in the core of the orogen, to large, open fold patterns and moderately to shallowly dipping fabrics in smaller structural domains in the north. keywords: geographical information systems, nagssugtoqidian orogen, west greenland, structural data, structural domains _______________________________________________________________________________________________________________________________________________________________________ j.a.m.v.g. & s.p., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: jvg@geus.dk s.p., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. present address: department of geolog y and geochemistry, stockholm university, 10691 stockholm, sweden. over the past ten years the nagssugtoqidian orogen in central west greenland (fig. 1) has been the subject of intense geological research, involving both bedrock mapping and research into the palaeoproterozoic and archaean tectonic evolution of the region. this has led to a major improvement in the understanding of the tectonic development of the nagssugtoqidian orogen (kalsbeek & nutman 1996; connelly et al. 2000; van gool et al. 2002). the research was undertaken in two projects, organised by the danish lithosphere centre (dlc) from 1994 to 1999 and the geological survey of denmark and greenland (geus) from 2000 to 2003, respectively. approximately 35 international geologists from institutions on three continents participated in the field work of these projects, with changing teams from year to year. during these projects a very large amount of data was collected, including structural measurements, lithological observations, intrusive relationships, information about metamorphic mineral assemblages, etc. other structural data were collected during previous work in part of the region in © geus, 2006. geological survey of denmark and greenland bulletin 11, 125–144. available at: www.geus.dk/publications/bull 126 68° 51° archaean, variably reworked metasedimentary rocks surficial deposits basalt quaternary palaeogene palaeoproterozoic sandstone sisimiut charnockite arfersiorfik quartz diorite granodioritic gneiss orthogneiss granitic gneiss dioritic gneiss anorthosite orthogneiss, undifferentiated amphibolite (includes proterozoic components) metasedimentary rocks (includes proterozoic components) (includes archaean components) metagabbro figs 2–8 qeqertarsuaq n n o c n o sn o nisb nordre isortoq nordre strømfjord ussu it nssz n ag ss ug to qi di an o ro ge n r in ki an fo ld b el t n or th a tla nt ic cr at on agto ataneq map sheet ussuit map sheet ilulissat disko bugt disko inland ice sydostbugtenaasiaat naternaq attu arfersiorfik kangerlussuaq itz sisimiut kangaatsiaq ikamiut jakobshavn isfjord 50 km inland ice greenland iceland canada fig. 1. schematic geological map of central west greenland (modified from van gool et al. 2002), depicting the nagssugtoqidian orogen. itz, ikertôq thrust zone; nisb, nordre isortoq steep belt; nssz, nordre strømfjord shear zone; sno, cno and nno, respectively, the southern, central and northern nagssugtoqidian orogen. small boxes outline the locations of the agto and ussuit map sheets; large box indicates the location of figs 2–8. 127 the 1960s and 1970s by the former geological survey of greenland (ggu, now part of geus). part of this was used by the collectors in their individual research or in small groups, and for the compilation of a published geological map, however, to date only a fraction of the total dataset has been made accessible in publications, and most of the original data collected prior to 2001 are only available for further analysis if extracted from the individual geologists’ field diaries and field maps. with traditional methods it would be highly impractical and tedious to obtain an overview of all the structural data collected over time from the nagssugtoqidian orogen in west greenland, and a rigorous conventional analysis of the complete dataset would be close to impossible. hence, alternative methods of data compilation and analysis were required. since 2001 geus’ mapping projects have included systematic collection of structural data which are available digitally. the data from all participants are recorded in field diaries and on maps, stored electronically in spreadsheets, and subsequently entered in a geographical information system (gis) for further presentation and analysis. in this way all data from an entire group of geologists can be accessed as a whole and used for map production and data analysis. gis methods has already proved useful in several disciplines including mineral exploration on local to global scales (e.g. bonham-carter et al. 1990; goodwin et al. 1996; knox-robinson & wyborn 1997; harris et al. 2001), palaeontology (e.g. carrasco & barnosky 2000), and environmental assessment (e.g. true et al. 1999; books 2000; wilson et al. 2000). in this paper we demonstrate the application of gis data management, visualisation and methods of analysis in a large-scale and long-term international project, including data from two previous mapping projects in the region. we present for the first time in a digital format a set of more than 10 000 structural orientation measurements and observations collected by more than 50 geologists in the nagssugtoqidian orogen over a period of 40 years. such a presentation can (a) provide a very helpful overview of the data itself, (b) help to identify where future research efforts may be scientifically interesting, and (c) show how these together with geological and geophysical maps can define structural domains and illustrate the large-scale structural variations through an important part of the orogen. the nagssugtoqidian orogen the nagssugtoqidian orogen in west greenland is a palaeoproterozoic collisional belt, dominated by archaean gneisses that were reworked at amphibolite and granulite facies during the palaeoproterozoic orogeny (van gool et al. 2002). it forms the northern boundary of the north atlantic craton in southern greenland, and is bound to the north by the contemporaneous rinkian fold belt. it consists of three tectonic segments, referred to as the southern, central and northern nagssugtoqidian orogen (sno, cno and nno; fig. 1), which respectively consist of a southern parautochthonous foreland, a high-grade core, and a transition zone to the rinkian fold belt. juvenile palaeoproterozoic magmatic arc rocks and supracrustal sequences occur mainly in narrow belts within the cno. the nagssugtoqidian orogen is characterised by a dominant ene–wsw structural trend, which culminates in a number of linear belts: the ikertôq thrust zone, the nordre isortoq steep belt, and the nordre strømfjord shear zone. these are interpreted as crustal-scale structures and alternate with areas dominated by large fold structures. detailed investigations in the core of the orogen have shown that during the nagssugtoqidian orogeny this region originally underwent a phase of nw-vergent thrusting, followed by folding now recognised predominantly as isoclinal folds (van gool et al. 2002). a second fold phase resulted in upright, ene-trending folds on a scale of tens of kilometres, with associated development of extension lineations plunging shallowly ene. finally, a phase of sinistral strike-slip shearing on the steep flanks of the large fold structures resulted in the above mentioned prominent linear belts. it is therefore only the latest deformation phases that generated the main ene–wsw-trending tectonic fabric of the orogen (van gool et al. 2002). the area discussed in this study extends from 67°n in the nordre strømfjord region to 69°10′n at jakobshavn isfjord and covers the northern part of the cno and most of the nno (fig. 1). structural data origin of the data the structural data have been derived from two different sources. the data north of 68°n were collected during recent geus mapping projects (2001–2003), while the data from south of 68°n have been extracted from published ggu and geus maps, collected during previous ggu, dlc and geus projects. during the recent geus 128 mapping projects, data were collected in the northern nagssugtoqidian orogen. structural measurements and other geological data were noted in field diaries together with their geographical coordinates using global positioning system (gps) receivers. these data were subsequently entered in spreadsheets and imported into arcview®. the geographical distribution of the structural data reflects the way they were collected along shorelines and on inland traverses. there may be several measurements at any one location, whereas no data were obtained in areas between traverses (which were often located many kilometres apart). in this study, we have restricted our analysis to the structural measurements. however, a combination of these data with other information, e.g. lithological and geophysical data, would make this gis-based analysis tool even more powerful. the southern part of the study area, south of 68°, is covered by two 1:100 000 scale maps, which were compiled prior to the digital storage of field data. the agto map sheet in the west (olesen 1984) consists of analogue data (but was recently digitised), whereas the ussuit map sheet in the east was produced in digital format (fig. 1; van gool & marker 2004). the structural data from the agto and ussuit map sheets are stored in geus’ geogreen map database and were extracted from this for the present study. however, these structural data only represent a fraction of the original data collected in the field. during the map compilations, the original structural data recorded on field maps or noted field diaries were filtered such that only representative measurements were shown on the final map; each measurement typically covers an area of a few square kilometres. thus, the southern part of the data compilation map in this paper shows an even distribution of data, and a much lower data density, compared to the more recently compiled areas in the north. the fact that the agto map sheet only contains very few lineation measurements compared to the surrounding regions also reflects a change in focus since the 1960s and 1970s, when measurement of lineations was not considered a high priority. a large gap in the data coverage occurs in the east, from arfersiorfik fjord to the north almost up to sydostbugten (fig. 1); this area was not covered by the mapping programmes by ggu and geus. definition of terms in the descriptions below the general orientation of a structural element is its three-dimensional orientation with respect to true north and horizontal, as defined by the combination of strike, dip direction and dip for planar structures, and plunge direction and plunge for linear structures. the term trend is used as the direction along the strike of a planar structure (without indication of dip direction), or the direction of plunge of a linear structure, without distinction between plunges up or down this direction. the trend is always expressed as two opposite directions (e.g. ne–sw). methods of data presentation for map presentation of the data we have used arcview® version 3.2. in addition, stereographic projections and statistical analysis for the determination of great circles and point maxima were prepared with stereonett (j. duyster, unpublished freeware). once the structural dataset has been incorporated into the gis database, the arcview® geoprocessing extension can be used to easily select subsets of data in areas with irregular shapes (fig. 2), or alternatively functions like arcview® query builder can be used to select data with certain characteristics. the data were plotted on a topographic map using conventional structural symbols, whereas orientations and dip angles were colour-coded. for foliations and lineations, four maps with different colour codes were plotted (figs 3–7). having attempted several different ways of displaying variations of dip/plunge directions on maps, we found 51°52°53° 50° 69° 68° 20 km 2 7 11 9 10 14 13 16 4 5 3 6 8 15 12 2 1 fig. 2. structural domains in the central and northern nagssugtoqidian orogen, based on structural orientation data. domain numbers refer to those used in table 1, fig. 8, and the main text. 129 that the regional trends were best shown using a subdivision into colour-coded quadrants. this gives a clear indication of variations on a regional scale and displays features that conventional plots of structural data would not have easily revealed. other features of the data could be highlighted with other methods of coding, or by plotting them on a different scale. the gis program allows the user to change the coding criteria and colours with a limited number of key strokes, and thus forms a powerful, user-friendly tool of analysis. we plotted one map for each of the planar and linear datasets (figs 3a, 5a), using differently coloured symbols for dip/plunge directions within each of four different quadrants: directions between 0–90° (ne quadrant) are shown in red, 90–180° (se quadrant) in orange, 180– 270° (sw quadrant) in green, and 270–360° (nw quad69° 51°52°53°54° 50° 68° 20 km foliation dip directions nenw sw se fig. 3a. foliation data, using four different colours to represent dip directions in the four quadrants of 0–90°, 90–180°, 180–270° and 270–360°. 130 rant) in blue. we also plotted foliations, lineations and fold axes, respectively, in three maps where the colour intensity reflects the steepness of the dip/plunge (figs 3b, 5b, 7). here the light orange colour indicates shallow dips/ plunges, and darker brown to black colours indicate progressively steeper dips/plunges. the data were also plotted in a third way by combining the two just described methods. the different colours were maintained for the dip/plunge directions within each of the four quadrants, combined with colour intensity to display the variations in dip. the foliation data were split into two separate plots to show more detail and avoid clutter: fig. 4a shows the overall ene–wsw-trending structures (blue and orange), whereas fig. 4b contains the overall ese–wnw-trending structures (red and green). a similar method was used for the lineations, however, on 69° 51°52°53°54° 50° 68° 20 km 0° < dip < 19° foliations 20° < dip < 29° 30° < dip < 49° 50° < dip < 69° 70° < dip < 90° fig. 3b. foliation data, using colour intensity to reflect dip angle. darker colours indicate steeper dip angles. 131 the scale of presentation many of the red and green lineations (ene–wsw-trending) would overlap. therefore, the green symbols were plotted separately (fig. 6a), whereas the red symbols were included with the blue and orange ones (fig. 6b); there are relatively few blue and orange symbols and therefore less cluttering. there are significantly less measurements of fold axes than of other structural elements. therefore we were not able to use their orientations for analysis of regional trends, and the fold axes are only colour coded for plunge angle (fig. 7). in some areas the very high data density causes a saturation with the colour of the main orientation on the scale of presentation. although the main trends can still be seen, minor orientation components may be obscured. this problem can be overcome by zooming in on smaller areas 69° 51°52°53°54° 50° 68° 20 km 0° < dip < 15° foliation azimuth and dip 16° < dip < 30° 31° < dip < 45° 46° < dip < 60° 61° < dip < 75° 76° < dip < 90° se nw sw fig. 4a. foliation data, using blue and orange colours for ne–sw trends with nw and se dips, respectively, combined with colour intensity to reflect dip angle. see fig. 4b for nw–se trends. 132 and printing on a different scale, revealing the full range of the data (e.g. mazur et al. 2006, this volume). structural domains apart from the general variations in structural style, it is apparent that there are well defined areas with distinct structural patterns. we therefore divided the whole study area into 16 structural domains (fig. 2), within each of which the structural characteristics are largely consistent and more or less distinct from those of adjacent domains. this subdivision is exclusively based on visual evaluation of the plotted data. a more rigorous approach for the definition of domains would have been possible, for example the method by vollmer (1990) based on eigenvalue cal69° 51°52°53°54° 50° 68° 20 km 0° < dip < 15° foliation azimuth and dip 16° < dip < 30° 31° < dip < 45° 46° < dip < 60° 61° < dip < 75° 76° < dip < 90° ne sw fig. 4b. foliation data, using red and green colours for nw–se trends with ne and sw dips, respectively, combined with colour intensity to reflect dip angle. see fig. 4a for ne–sw trends. 133 culations of data in small subsets, but is beyond the scope of the present study. the domains are presented schematically in fig. 2, and their outlines are also shown in figs 3–8. the structural data for each of the domains were extracted and plotted as equal area, lower hemisphere stereographic projections (fig. 8). foliations were plotted as poles to planes and contoured, and the orientation of the maximum density of data indicated in each plot. great circles were calculated where visual inspection of the contoured data suggested that a great circle distribution exists. calculations of great circle and fold axis orientations are based on the orientations of the three eigenvectors of the data. the large number of data in the contoured plots results in an accentuation of the high concentrations, while smaller populations that define separate structures are less visible. however, these are included in the calculations of 69° 51°52°53°54° 50° 68° 20 km lineation orientations nenw sw se fig. 5a. lineation data, using four different colours to represent plunge directions in the four quadrants of 0–90°, 90–180°, 180–270° and 270– 360°. see also fig. 5b. 134 the great circles and fold axes, and therefore the calculated great circle may diverge from the one defined by the maximum orientations, as in domains 7, 9 and 14 (fig. 8). lineations were plotted and contoured, with indication of the orientation of the maximum concentration of data (fig. 8b). table 1 contains short descriptions of the characteristics of each domain regarding foliation, linear data and general geology. results the main variations in the structural patterns within the study area are described in the following sections, using structural maps and stereographic projections (figs 3–8). the structural variations are apparent at a first glance as clustering of data and variations in colours; they reflect the nature of the large-scale tectonic evolution of the orogen, which is discussed in a final section. 69° 51°52°53°54° 50° 68° 20 km 0° < dip < 15° lineations 16° < dip < 30° 31° < dip < 45° 46° < dip < 60° 61° < dip < 90° fig. 5b. lineation data, using colour intensity to reflect steepness of plunge. see also fig. 5a. 135 foliations overviews of the orientations of the planar fabrics are shown in figs 3, 4. the predominant foliation trend is ene–wsw, shown in orange and blue colours. linear belts in this direction, dominated by steeply dipping foliations, alternate with broader regions characterised by strongly variable orientations. these belts and regions with different structural characteristics have previously been referred to as steep belts and flat belts, respectively (marker et al. 1995). the distinct alternation between such distinct linear belts and folded regions diminishes towards the north, and the predominant general ene–wsw trend becomes progressively weaker, as reflected by the increase of redand green-coloured symbols. this is apparent especially in the north-eastern corner of the study region, where the foliations are dominated by ne dip directions 69° 51°52°53°54° 50° 68° 20 km 0° < plunge < 15° lineation direction and plunge 16° < plunge < 30° 31° < plunge < 45° 46° < plunge < 60° 61° < plunge < 75° 76° < plunge < 90° sw fig. 6a. lineation data, using green colour to represent sw directions and colour intensity to reflect steepness of plunge. see fig. 6b for nw, ne and se directions. 136 shown in red. coupled with this progressive change towards the north there is an overall decrease in the dip angle, as expressed by an increasing amount of light orange-coloured symbols. in the south, the two main linear belts, the nordre isortoq steep belt in domain 1 and the nordre strømfjord shear zone in domain 4, are characterised by a near-uniform ene–wsw-trending foliation, a marked absence of nw–se-trending foliations, and steep dip angles. a third linear belt in the north, the naternaq belt in domains 12 and 13 and the northern part of domain 7, is discontinuous and less well defined. smaller, discontinuous shear zones also occur in the nno e.g. in domains 10 and 15; these are indicated by strong clustering and alignment of symbols of the same colour, but not necessarily by steep dips. dip directions of the overall ene–wsw-trending 69° 51°52°53°54° 50° 68° 20 km decreasing colour intensity = decreasing angle of plunge lineation direction and plunge sw se nw ne fig. 6b. lineation data, using blue, red and orange colours for nw, ne and se orientations, respectively, combined with colour intensity to reflect steepness of plunge. see fig. 6a for sw directions. 137 structures (in blue and orange colours, fig. 3a) show a clear regional pattern of alternating nw and se dip directions. in the south, switches in dip directions are associated with the two main linear belts such that the intervening area, which forms a large anticlinorium (van gool et al. 2002), is characterised by predominant sse dips (in orange), while nnw dips (in blue) prevail to the north and south. nw–se-trending foliations are predominant in two distinct areas in the north-east: one in the extreme north-eastern corner with predominating ne-dipping foliations (in red), and another around sydostbugten, characterised by sw-dipping foliations (in green). farther south only two areas of uniform dip directions are recognised, one around attu with predominant sw dips (in green), and another forming a belt north of the nordre strømfjord shear zone, which has uniform nw dips (in blue). 69° 51°52°53°54° 50° 68° 20 km 0° < plunge < 15° fold axes 16° < plunge < 30° 31° < plunge < 45° 46° < plunge < 60° 61° < plunge < 75° 76° < plunge < 90° fig. 7. fold axis data, using colour intensity to reflect steepness of plunge. 138 69° 51°52°53°54° 50° 68° 20 km n = 233 max. density at 337/78 2 3 4 6 7 8 11 9 10 15 16 1 12 13 14 n = 231 max. density at 337/78 n = 529 max. density at 346/78 5n = 630 max. density at 347/54 n = 1239 max. density at 323/48 fa 268/27 n = 746 max. density at 335/72 fa 244/26 n=198 max. density at 175/78 n = 829 max. density at 330/24 fa 246/4 n = 343 max. density at 163/42 fa 244/4 n = 569 max. density at 260/18 fa 251/14 n = 384 max. density at 128/48 n = 476 max. density at 153/66 fa 72/33 n = 552 max. density at 180/24 fa 251/6 n = 743 max. density at 129/42 fa 50/20 n = 1152 max. density at 60/24 fa 67/21 n = 1238 max. density at 171/84 fa 80/24 2 1 2 3 4 5 12 6 7 8 1110 foliations 14 13 15 16 5 4 9 fig. 8a. characterisation of structural domains from fig. 2 with stereographic projections of foliations within each domain. poles to foliations plotted on lower hemisphere, equal angle nets and contoured at 1, 2, 3, etc. times random distribution. the number of data points (n) and orientation of maximum density are indicated for each plot. fa, calculated fold axis. 139 69° 51°52°53°54° 50° 68° 20 km 2 3 4 5 6 7 8 11 9 10 15 16 1 12 13 14 n = 27 max. density at 60/18 n = 309 max. density at 78/12n = 29 max. density at 72/30 n = 86 max. density at 253/06 n = 59 max. density at 49/24 n = 308 max. density at 267/24 n = 547 max. density at 240/12 n = 92 max. density at 267/24 n = 592 max. density at 244/06 n = 288 max. density at 240/12 n = 224 max. density at 254/12 n = 139 max. density at 129/48 n = 217 max. density at 72/0 n = 465 max. density at 258/12 n = 373 max. density at 60/24 n = 525 max. density at 45/18 2 1 2 3 4 5 6 7 1110 lineations 14 13 15 16 5 4 9 12 8 fig. 8b. characterisation of structural domains from fig. 2 with stereographic projections of lineations within each domain. data plotted on lower hemisphere, equal angle nets and contoured at 1, 2, 3, etc. times random distribution. the number of data points (n) and orientation of maximum density are indicated for each plot. 140 steep n-dipping foliation with consistent orientation. a sharp transition to s-dipping foliations at the northern boundary of the belt. foliation curved along shallowly ene-plunging folds on a scale of tens of kilometres. dips variable, se dips predominating over ne and less common sw dips. steep foliation. both northerly and southerly dips. few folds on scales of 0.5–1 km. steep foliations with consistent orientations, slightly oblique to the trend of the linear belt. northerly dips predominate. near the southern boundary a sharp transition to sse-dipping foliations. gradual northward transition from steep dips close to the nssz to shallower nnw dips. the foliation locally curves into the shear zone. ne dips are mixed with the dominant nnw dips in most of the area. the plentiful data points on the stereonet (fig. 8a)obscure orientations (mainly ne-dipping) away from the maximum. intensely folded region with 5–10 km large folds, mainly w-plunging. northerly and westerly dips predominate in contrast to domain 5. variable strike. nnw dips less predominant than in domain 5. folding. dips mainly moderate, but steep in the south-west. fig. 8a displays a girdle over shallow, wsw-plunging fold structures. a calculated great circle is discordant to the measured maxima, which align on a steeper great circle. foliations outline the large fold structure visible on fig. 1. southerly dips predominate in the northern part (green and orange/brown, fig. 3a). a zone with nw dips (blue) runs through the domain centre. mixed dip directions in the south. dips moderate to steep. irregular stereonet pattern (fig. 8a), with remnants of a great circle distribution similar to that in domains 6 and 7. foliation lineation geology synopsisdomain 1 2 3 4 5 6 7 8 consistent, predominant shallow ene plunge. variable lineations, with predominant shallow ene plunges, parallel with fold axes. shallow, se-plunging lineations in a c. 10 km wide zone south of the nssz. lineations mainly shallowly ene-plunging. some variation in fold hinges, especially near the eastern domain boundary. subhorizontal lineations, the majority wsw-plunging in contrast with surrounding areas. progressive change from shallow ene plunges near the nssz towards steep ne and n plunges in the north-west. mixed orientations with a cluster of shallowly se-plunging lineations. in the west progressive change in orientation continues from domain 5. more variable orientations in the east. rather consistent enetrending lineations. ene plunges in the east become shallower and mixed with wsw-plunging domains in the west, followed by moderate to steep wsw plunges at the coast. locally steep lineations in fold cores. steep swand se-plunging lineations around a large central fold core. mainly e–w-trending lineations in the west in transition to domain 7. variable lineations in the east. two maxima of shallow w plunges and steeper se plunges (fig. 8b), the latter possibly with a small circle distribution. nordre isortoq steep belt (shear zone), predominant sinistral shear. mainly paragneiss. northern cno flat belt of interleaved archaean orthogneisses and proterozoic orthoand paragneisses. steep belt within the northern cno flat belt; tightly interleaved orthoand paragneisses. nssz. interleaved orthoand paragneisses. archaean orthogneisses and horizons of supracrustal amphibolite. archaean indentor, orthoand paragneiss as in domain 5. large fold structures. archaean orthogneisses and metasedimentary rocks define a poorly sampled linear zone at the northern boundary of the indentor block of piazolo et al. (2004). a few kilometre-sized fold structures. orthogneisses interleaved with supracrustal amphibolite. a fold interference pattern occurs south of the main fold at the western end of the naternaq belt. high-strain strike-slip zone constrained by metasedimentary rocks. anticlinorium between two shear zones. lineation constant in spite of intense folding. only one zone where lineations plunge se, locally steeply. steep zone with tight folds within the larger domain 2 with large fold structures. steep zone discontinuous to the east, and concentrated in an area dominated by two metasedimentary belts. sinistral strike slip zone5 km wide. obliquity of foliation fits with sinistral shear. region with shallow, n-dipping structures. overall wand wsw-plunging folds. two sets of lineations, one variable, the other consistently wsw-plunging, parallel with fold axis. predominantly straight and steep foliations with ene trends, but variable dip directions, and consistently shallow lineations. misalignment of foliation girdle and calculated great circle (fig. 8a) indicates complex fold pattern: along-strike variation of lineations (and presumably fold axes) and local folds with steep axes disturb the stereographic plot. non-consistent dip directions and dips in the northern linear zone suggest intense folding. mixed structural patterns including ene-trending fold limbs and fold interference patterns. the core of the large fold with the steep lineations is located along the southern extension of a nne-trending belt of steep lineations on the western limb of the naternaq supracrustal belt. table 1. summary of characteristic features of the structural domains cno: central nagssugtoqidian orogen. nno: northern nagssugtoqidian orogen. nssz: nordre strømfjord shear zone. 141 consistent ene trend with both nnw and sse dip directions. large fold structures in the east. steep dips, shallower towards north. great circle distribution with shallowly nnwdipping and subordinate steep, ssw-dipping flanks tentatively interpreted as due to asymmetric s-vergent folds (fig. 8a). mainly steep s dips with consistent trend. as in domain 9, ene-wswtrending foliation in the west swings towards e–w in the east. at the southern boundary a linear belt with moderate s dip. partial great circle distribution; no dip directions within the ne quadrant (fig. 8a). foliations outline a w-plunging antiform c. 10 km large. the northern part of its northern limb appears overturned towards n. great circle distribution with predominant shallow to moderate dips on stereonet (fig. 8a). foliations outline a fold with a folded, overall moderately se-dipping western main limb. the high-strain southern limb dips steeply s. large spread on stereonet (fig. 8a). eastern continuation of steep, high-strain southern limb of fold from domain 12 and large e-plunging antiform. nw dips more common in the north, dips shallower near hinge. the northern limb has moderate to steep ne dip. fig. 8a shows a well-defined great circle and ne-plunging calculated fold axis. irregular foliation in the core of the eastern naternaq fold, forming an e–w-trending whaleback structure. fig. 8a shows a point maximum and partial girdle which do not fit the calculated great circle (see the main text). foliations define a large, open ne-trending synform with steepest orientations in the core, bounded by straight belts. an antiform occurs in the south-east, with its southern limb in domain 13. a well-defined great circle on fig. 8a indicates cylindrical, ne-plunging folds. the foliation defines large, open, ne-plunging folds besides a large antiform surrounding the synform of domain 15. the stereonet data (fig. 8a) display a point maximum with a partial great circle distribution. foliation lineation geology synopsisdomain 9 10 11 12 13 14 15 16 predominant subhorizontal ene–wsw-trending lineations, gradually changing to e–w in the east, where large-scale folds occur. very strong preferred ene–wsw subhorizontal orientation, with a tail towards e–w trends displayed on fig. 8b. consistent ene–wsw-trending subhorizontal lineations, with indistinct domains of respectively easterly and westerly plunges. strong point maximum on stereonet with shallow wsw plunges (fig. 8b). shallow plunges of lineations, mainly towards w. subordinate nw plunges on northern fold limb. fig. 8b shows strong point maximum parallel with calculated fold axis. moderately se-plunging lineations on the western limb, with isolated ssw plunges in hinge zone. very few measurements on the southern limb, with shallow wsw plunges. lineations shallow and eneor wsw-plunging on the southern limb. variable plunges on the eastern limb between ne and ese. the latter orientation most common in hinge areas. fig. 8b clearly shows these three separate populations. predominant subhorizontal wsw plunges, except ene plunges at the eastern domain margin along the eastern limb of the map-scale fold. fig. 8b shows a single strong point maximum. strong predominance of ene-plunging lineations along the synform axis. consistent se plunges south-east of the shear zone. consistent moderate nw plunge in the antiform near head of fjord. fig. 8b shows a fairly well defined point maximum close to the calculated fold axis. predominant shallow, ne-plunging lineations and a small population of slightly steeper, e-plunging lineations in the north-eastern domain corner. the lineations swing, following the folds. region around kangaatsiaq with metasedimentary rocks, amphibolite and granite within the regional grey gneiss. straight zone of predominant orthogneisses around aasiaat, bounded to the south by a high-strain zone. ikamiut supracrustal rocks. large antiform with 1 km-scale parasitic folds on its southern limb. northern limb poorly exposed and undersampled. western naternaq belt. steep paragneisses folded on 20 km-scale. mainly data from isoclinal fold on western limb; less from straight southern limb. steeply dipping, straight gneisses in eastern naternaq belt. a large e-plunging antiform at the eastern end, and a north-eastern fold limb with only minor metasedimentary rocks. predominantly orthogneiss, cut by flat-lying shear zones. synform with strongly migmatitic paragneiss and a core of migmatitic orthogneiss. high-strain zones to the north-west and south-east. archaean orthogneiss with thin sheets of pelitic rocks and supracrustal amphibolite. also areas of very weakly foliated porphyritic granodiorite. transition from predominant steep, northerly dips in the cno to shallower, variable dips and large fold structures in the north. from east to west a large swing from ene–ssw to e–w trends. lineations uniform also in the area of large folds. very consistent foliation trends, the northernmost widespread steeply dipping foliations, and very persistent lineations. data consistent with a shallowly w-plunging antiform becoming progressively tighter westward (but not easily traced into domain 10). western limb of naternaq supracrustal belt, forming a distinct zone, apparently transecting the overall ene-trending fabric, and with uncommon se-plunging lineations and fold axes. consistent lineation trend in antiform, but opposite plunge directions on the limbs. steeper lineations in the fold hinge. very consistent lineation trends, also through the antiform in the east. direction of plunge flips over in the east, perhaps indicating two generations of lineations. the south-eastern limb of the synform, overturned to the north-west and becoming very tight towards north-east. distinctly different lineations in the underlying, folded shear zone. the ne-dipping orientations and open folds are significantly different from elsewhere in nno. the shear zone exposed on either side of the synform in domain 15 does not continue in domain 16. table 1 (continued) 142 the stereographic projections of the foliations clearly reflect two trends (fig. 8a). the plots from the southern part of the study area define alternating point maxima and great circle distributions, reflecting, respectively, the linear belts and the fold-dominated regions. in contrast, in most of the nno the plots mainly display (partial) great circle distributions or otherwise irregular patterns, indicating the lack of extensive linear belts in the north. furthermore, the stereographic projections reflect the northward decrease in dip angle: in the cno the point maxima indicate dip angles around 80°, whereas the nno is characterised by point maxima indicating dip angles in the range 20–40°. lineations overviews of the linear fabrics are shown in figs 5, 6. the highly uneven data density in the lineation maps (fig. 5a, b) reflects that the southern and northern parts of these maps have been compiled from different sources, i.e. published maps in the south and complete field datasets in the north. nevertheless, it is apparent that the dominant trend is ene–wsw, as indicated by the predominant red and green colours (fig. 5). outside the linear belts, gradual changes in the lineation trends on a scale of 10–50 km or more are seen, for example between attu and nordre strømfjord, at western ussuit, and in the north-east of the study area. several smaller areas are dominated by e–w trends, for example east of attu, south-east of sydostbugten, and south of jakobshavn isfjord. most lineations plunge 0–30°, with markedly steeper plunges in regions of map-scale fold interference structures and fold hinges. this is prominent west of ataneq, north-west of attu, and in the eastern part of domain 13. these regions are also characterised by orientations that diverge from the general ene–wsw trend. the structural maps (fig. 6) stereographic projections (fig. 8b) clearly show that wsw-plunging lineations predominate in most of the nno (north of the nordre strømfjord shear zone) except for small clusters of neplunging lineations and the area east and north-east of sydostbugten, where the plunge is towards ene (domains 13, 15 and 16); there is a sharp break between these two plunge directions east of sydostbugten. the same pattern is shown by the fold axes calculated from the great circle girdles of the foliations (fig. 8a). south of the nordre strømfjord shear zone (in the cno), the point maxima of the lineations and calculated fold axes consistently indicate shallow ene plunges. fold axes orientations of fold axes are shown on fig. 7. the symbols are colour-coded for plunge angle in order to facilitate comparison with the lineation data. the orientations of the fold axes mimic the general characteristics of the lineations, being generally subparallel with the latter. their distribution in clusters reflects a higher density of measurements in areas of map-scale fold hinges, where outcrop-scale folds are more common. tectonic implications the structural data presented in this paper show that the tectonic style changes significantly across the central and northern parts of the nagssugtoqidian orogen. while the cno is dominated by steep, linear and continuous belts separated by well-defined areas of large-scale folding, the nno does not contain such continuous, linear belts, whereas 20–80 km-scale folds are abundant. the main change in tectonic style occurs across the nordre strømfjord shear zone. more specifically, the southern part of the study area (the cno and the nordre strømfjord shear zone) is dominated by alternating linear belts and folded regions. the corresponding structural domains follow the main strike of the linear belts and are continuous from the coast to the inland ice. the linear belts themselves are dominated by sinistral strike-slip deformation. in contrast, the domains in the nno are generally less elongate and reflect the lack of linear belts of similar dimensions as in the cno. small high-strain zones are observed locally, e.g. along the northern and southern margins of domain 15 and with several examples in domains 7, 9 and 10. shear sense indicators are rare and inconsistent in the nno, and the overall deformation in this region seems to be predominantly coaxial (piazolo et al. 2004; mazur et al. 2006, this volume). the change in tectonic style is interpreted to be a result of (a) differences in localisation of strain as high-strain, steep belts, (b) different deformation kinematics (strike-slip wrench tectonics in the south, coaxial deformation in the north), and (c) variations in the intensity of deformation. we consider that the overall palaeoproterozoic strain is significantly lower in the nno than in the cno. the nno commonly preserves shallow dips, which presumably predominated after the original phase of thrusting. besides, the two latest deformation phases, which are responsible for the steep foliation in the cno, are less intense in the nno (van gool et al. 2002; piazolo et al. 2004; mazur et al. 2006, this volume). these observations may account for the previously 143 outlined differences in the mode of strain localisation. the data presented here furthermore show that the change in style is rather abrupt across the nordre strømfjord shear zone, and thus support the interpretation by sørensen (1983) and sørensen et al. (2006, this volume) that this structure is of crustal scale and has a significant offset – a notion that has previously been questioned by hanmer et al. (1997). it is well established in the literature that the structural pattern of the cno is fully attributed to nagssugtoqidian deformation (van gool et al. 2002). in contrast, the nno is currently interpreted as having only in part been affected by palaeoproterozoic deformation, and the nagssugtoqidian strain is furthermore partitioned into smaller regions (piazolo et al. 2004; mazur et al. 2006, this volume). a significant part of the deformation in the nno thus seems to be of archaean age, and its overall structural style defined by interference between archaean and palaeoproterozoic structures. therefore, like mazur et al. (2006, this volume) we suggest that the change in tectonic style from south to north reflects partitioning of nagssugtoqidian strain. this is clearly illustrated by the smaller and less elongate structural domains in the nno, and by the significant change of the general trend of both foliations and lineations towards the north-eastern corner of the study area, where we consider that the influence of the palaeoproterozoic deformation diminishes rapidly. this interpretation is supported by the relatively low metamorphic temperatures recorded by hollis et al. (2006, this volume) in some parts of the nno. it is beyond the scope of this contribution to explore the details of the structural domains that we have outlined. however, three other contributions in the present volume of geological survey of denmark and greenland bulletin deal with the specific nature of some of these domains. sørensen et al. (2006, this volume) investigate the character of the nordre strømfjord shear zone and adjacent areas in domains 4 and 5, mazur et al. (2006, this volume) focus on the partioning of structures within domains 6, 7 and 8, and hollis et al. (2006, this volume) describe structures within domains 11 and 15. conclusions the application of a gis computer program enables us to visualise large amounts of structural data in a variety of ways. thus, we can rapidly obtain an overview of large datasets that are otherwise difficult to manage, and delineate areas with consistent tectonic trends. although the methods applied here de not reveal features that are not present in the original geological maps, they can substantially facilitate the detection and description of structural trends and variations. in addition, stereographic plots of each of the domains can quickly be produced and analysed. in the present case, the methods greatly helped to subdivide the central and northern nagssugtoqidian orogen into distinct structural domains with specific individual characters. the investigation of the large-scale structural trends in the central and northern nagssugtoqidian orogen revealed distinct changes in the tectonic style from south to north. in the core of the orogen, the ene-striking nordre strømfjord shear zone from the coast to the inland ice forms the most prominent feature. the tectonic character of the orogen changes across this structure from predominantly ene-tending, steep fabrics in the south to large fold patterns and generally flat structures in the north, with a marked decrease in the main dip angle from c. 80° south of the shear zone, to c. 20–40° in the north. in addition, a significant decrease in the intensity of deformation is apparent, coupled with a decreasing proportion of the strain localised in linear belts. we interpret these patterns as reflecting a general northward decrease in the palaeoproterozoic tectonic overprint on archaean structures, as well as strain partitioning in smaller regions. hence, some of the structural domains in the nno are largely unaffected by pervasive palaeoproterozoic deformation. acknowledgements reviews by john grocott and ken mccaffrey are gratefully acknowledged. references bonham-carter, g.f., agterberg, f.p. & wright. d.f. 1990: weights of evidence modelling: a new approach to mapping mineral potential. geological survey of canada paper 89, 171–183. books, c.j. 2000: defining groundwater system recharge and vulnerability areas in regions of suburban expansion; overview of the northern illinois example. abstracts with programs – geological society of america 33, 45 only. carrasco, m.a. & barnosky, a.d. 2000: miomap: a gis-linked database to assess the effects of tectonic and climatic changes on mammalian evolution. abstracts with programs – geological society of america 32, 15 only. connelly, j.n., van gool, j.a.m. & mengel f.c. 2000. temporal evolution of a deeply eroded orogen: the nagssugtoqidian orogen, west greenland. canadian journal of earth sciences 37, 1121–1142. goodwin, p.b, choiniere, m.e., harris, f.w. & dean, b.p. 1996: im144 proving exploration with geographical information system (gis) technology. american association of petroleum geologists bulletin 80, 1297. hanmer, s., mengel, f., connelly, j. & van gool, j.[a.m.] 1997: significance of crustal-scale shear zones and syn-kinematic mafic dykes in the nagssugtoqidian orogen, sw greenland: a re-examination. journal of structural geology 19, 59–75. harris, j.r., wilkinson, k., heather, k., fumerton, s., bernier, m.a., ayer, j. & dahn, r. 2001: application of gis processing techniques for producing mineral prospectivity maps – a case study: mesothermal au in the swayze greenstone belt, ontario, canada. natural resources research 10, 91–124. hollis, j.a., keiding, m., stensgaard, b.m., van gool, j.a.m. & garde, a.a. 2006: evolution of neoarchaean supracrustal belts at the northern margin of the north atlantic craton, west greenland. in: garde, a.a. & kalsbeek, f. 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revision accepted 20 december 2005 geological survey of denmark and greenland bulletin 17, 2009, 81-84 the evolution of the landscape of north-east brazil in relation to the burial and exhumation history of both onshore and offshore areas is the focus of a research project carried out for statoilhydro do brasil and petrobras from 2007 to 2009 by the geological survey of denmark and greenland in collaboration with geotrack international. in hydrocarbon exploration it is important to understand the regional tectonic framework and thus also to consider the volumes of rocks that may have been present and then removed during the geological past. for example, the timing of hydrocarbon generation and changes in migration routes can be assessed when the timing and magnitude of uplift and erosion is known. studies in west greenland have demonstrated the usefulness of large-scale, low-relief, high-level landscapes as markers of uplift events, and in particular the strength of combining the denudation history from landscape analysis with the cooling history from apatite fission-track analysis (afta) data and the stratigraphic record (bonow et al. 2006, 2007; japsen et al. 2006, 2009). in the study area, there are two plateaux with elevations up to c. 1300 m above sea level (a.s.l.). the plateaux are currently being dissected by deeply incised fluvial valleys, and escarpments separate the two plateaux. the lowlands cut across early cretaceous rift systems along the atlantic margin, including the intracontinental recôncavo–tucano–jatobá (rtj) rift and also the camamu basin, of which the western margin is exposed onshore (fig. 1). the rtj rift is a mature hydrocarbon province, whereas the deep-water parts of the camamu basin are the target of frontier exploration (e.g. magnavita et al. 1994; davison 1999; cobbold et al. 2008). the post-rift sequence in the rtj rift and the inshore camamu basin is thin or absent. however, it has been estimated that up to 2000 m of sedimentary cover once was present, but has now been removed (magnavita et al. 1994). the atlantic margin of brazil is characterised by elevated plateaux cut by deeply incised valleys, but this landscape has a pattern common with many other passive continental margins with elevations from 1000 to 2000 m a.s.l. or more around the world, for example in norway, east and west greenland and south-east australia. mesozoic–cenozoic rift systems parallel to the coast are generally present offshore 81 post-rift landscape development of north-east brazil johan m. bonow, peter japsen, paul f. green, peter r. cobbold, augusto j. pedreira, ragnhild lilletveit and dario chiossi © geus, 2009. geological survey of denmark and greenland bulletin 17, 81–84. available at: www.geus.dk/publications/bull ~100 km m a.s.l. fig. 2a fig. 2b fig. 3 brazil 225 km 42°w 10°s 12°s 14°s 40°w rtj 38°w neogene deposits cenozoic laterites barreiras fm (neogene) sabiá fm (lower miocene) mesozoic deposits palaeozoic deposits precambrian rocks a bbb chapada diamantina ppppppppppppppppppppppppppppppppppppppppppppppppppppppppppppppppppppppppppppppppppppppppppppppppp lllllllllllllllllllllllllllllllllllllllllllllllllllllllll aaaaaaaaaaaaaaaaaaaaaaaaa nnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnn aaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaa llllllllllllllllllll ttttttttttttttttttttt oooooooooooooooooooooo dddddddddddddddddddddd eeeeeeeeeeeeeeeeeeeeeeeeeeeeeeeeeee ccccccccccccccccccccccccccccccccccccccccccccccccccccccccccc oooooooooooooooooooooooooooooooooooooooooooooooooooooooooooooooooooooo nnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnnn qqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqqq uuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuuu iiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiii ssssssssssssssssssssssssssssssssssssssssssssssssssssssssssssssssssssssssssssssssss tttttttttttttttttttttttttttttttttttttttttttttttttttttttttttttttttttttttttttttttttttttttttttttttttttttttttt aaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaaa p l a n a l t o d e c o n q u i s t a camamu basin 2000 1800 1600 1400 1200 1000 800 600 400 200 0 10°s 12°s 14°s 16°s 42°w 40°w 38°w fig. 1. a: geological map of the study area (based on cprm 2001, 2003). precambrian basement is covered by younger sedimentary sequences, which are important age constraints for the different peneplains. rtj, recôncavo – tucano–jatobá basin. the arrow points at the small sabiá formation outrcrop. b: topography of the study area. two topographical features dominate the landscape: a lower surface, which is a plain mainly at 200–500 m a.s.l. (greenish and light yellow) and the higher surface which is a plain mainly at 900–1200 m a.s.l. (orange and reddish). pronounced escarpments separate the two features. elevation data source: jarvis et al. (2008). rosa_2008:rosa-2008 01/07/09 15:49 side 81 with a transition from continental to oceanic crust farther offshore. several aspects related to elevated, passive continental margins are controversial: the origin of the plateaux, the timing of their uplift to their present elevation and their relation to the adjacent rift systems (japsen et al. 2009). for example, gallagher et al . (1994) found that the almost 3 km high mountain chain serra do mar, near rio de janeiro, is the remnant of a rift shoulder from the early cretaceous breakup in the south atlantic, whereas cobbold et al. (2001) argued that these mountains were formed during neogene block-fault tilting. king (1967) mapped stepped surfaces (i.e. erosion surfaces at distinct levels in the landscape) through the elevated terrains along many passive continental margins, e.g. in eastern australia, southern africa and north-east brazil. king used remnants of sedimentary rocks to constrain the ages of these surfaces and concluded that the main surfaces were formed during the cenozoic, and consequently that the margins had been uplifted in that same time interval. further more, king found that the highest peaks in the interior of the continents represented remnants of a pre-break-up topography. subsequent geomorphological research into the development of the passive margins of southern africa and eastern australia has, however, regarded the elevated terrains along these passive margins as mainly reflecting preserved riftshoulders (e.g. ollier 1985), a model that is commonly used as input to thermochronological studies (e.g. gallagher et al. 1994). the burial and exhumation history of the study area is currently investigated by combining the cooling history from apatite fission-track analysis from both outcrop and borehole samples with the denudation history from landform analysis and the stratigraphic record. based on field work carried out during four weeks in july and august 2007, we here report results focused on the rift systems near the atlantic margin and on the interior highlands in north-east brazil. large-scale landforms a geomorphological analysis based on the method described by bonow et al. (2006) has led to the identification of two major erosion surfaces (peneplains) of low, relative relief and of regional extent within the study area: a lower surface extending from near-coast areas and far into the interior (up to 500 m a.s.l., greenish colours in fig. 1b) and a higher surface that includes the plateaux (‘planaltos’) of chapada dia mantina (c. 1200 m a.s.l., reddish colours in fig. 1b) and planalto de conquista (fig. 1b; c. 900 m a.s.l.). both surfaces cut across rocks of different ages and resistances and must therefore be erosional features (fig. 2). these observations show that the surfaces were originally formed by denudation to a near-horizontal plain because even slightly tilted surfaces will be incised by fluvial valleys and the relief rejuvenated, until a new and younger peneplain is formed. the location of these surfaces near the present coast indicates that sea level was the most likely baselevel to which the surfaces graded. after uplift of what is now the higher surface, the lower surface was developed by incision along the main rivers in the area (figs 1, 3, 4). the higher surface can be correlated from chapada dia mantina to planalto de conquista at a slightly lower elevation (fig. 2). it is characterised by an undulating plain with shallow and wide valleys, and it is preserved on high ground with resistant rocks. the higher surface must have developed across a larger area than that presently preserved, as the lower surface has developed at the expense of it. escarpments often 82 w w e e ? ? 0 1 2 0 1 2 precambrian basement jurassic–cretaceous sediments maximum topography in corridor minimum topography in corridor higher surface lower surface a b 40°w 39°w 38°w 42°w 41°w 40°w 39°w km a .s .l. km a .s .l. fig. 2. two profiles to illustrate the surface mapping. the dotted lines show the interpreted peneplains. the surfaces cut across both the basement, consisting of rocks that have different resistance to erosion, as well as the sedimentary sequence of the rtj basin. this shows that the peneplains are erosional features. see fig. 1b for profile location and corridor width. rosa_2008:rosa-2008 01/07/09 15:49 side 82 separate the lower surface from the higher surface (figs 3, 4). in detail, these escarpments usually coincide with bedrock boundaries, thus reflecting bedrock resistance. the valley patterns and the incision of rivers in the recôncavo and tucano basins also show that the lower surface is rapidly being dissected by incising rivers, due to a change in baselevel after formation of the lower surface. geological constraints the lower surface cuts across post-rift strata within the rift and precambrian basement outside the rift (figs 1, 2). the formation of the surface thus post-dates the aptian marizal formation (e.g. magnavita et al. 1994). the age of the surface may, however, be further constrained by an outlier of the early miocene, marine sabiá formation within the recôn cavo basin (fig. 1a) where this sedimentary unit has been found in deep trenches (viana et al. 1971). this outlier testifies to a marine transgression that occurred before the formation of the lower surface because the outlier is truncated by that surface, and thus the lower surface is younger than early miocene. the areas where the higher surface is defined are characterised by laterites (deep weathering profiles) of cenozoic age (fig. 1; cprm 2001, 2003). both the higher surface and the laterites are currently being destroyed by erosion along the escarpments that outline the plateaux. consequently, the laterites must have formed at the end of the erosional process that shaped the higher surface. based on the age of the laterites we can deduce that the higher surface formed during the cenozoic. this time interval can be further narrowed if we take into account that the younger, lower surface was formed subsequent to the deposition of the sabiá formation. the age of the higher surface may thus tentatively be estimated to be palaeogene, which implies that the landscape at that time was a peneplain close to sea level. this suggestion is in agreement with observations from similar plateaux north and south of the study area where the plateau surfaces were exposed during the palaeogene according to stratigraphic data (sant’anna et al. 1997; morais neto et al. in press) and geochronological constraints on deep weathering (spier et al. 2006; lima 2008). alternatively, both the higher and the lower surface in the study area may have formed during the neogene, which also agrees with the cenozoic age of the laterites. geo morphological analysis alone cannot definitely conclude which alternative is correct, but we prefer the first alternative because it is consistent with independent constraints from outside the study area. 83 fig. 3. digital terrain model with a higher surface (hs) and a lower surface (ls) that are separated by escarpments in the chapada diamantina area. the higher surface forms a coherent plateau at 1200–1400 m a.s.l. (reddish) with only minor valley incisions. this plateau is presently being dissected by rivers along the escarpment, erod ing down to the lower surface, here at 500–400 m a.s.l. (greenish). escarpments are also found above the higher surface, maybe representing steps towards older surfaces at higher elevations. the arrow indicates the location and direction of the photograph in fig 4. elevation data source: jarvis et al. (2008). fig. 4. the lower surface with the escarpment and the higher surface in the background. see fig. 3 for location. 2000 1500 1000 500 m a.s.l. hs ls 42°w 41.5°w 41°w 14 °s 13 .5° s 13 °s ~25 km rosa_2008:rosa-2008 01/07/09 15:49 side 83 conclusions the landscape in the study area is dominated by two main peneplains, a higher surface and a lower surface that were both formed as low-relief erosion surfaces. the higher surface developed during the cenozoic, probably during the palaeogene as other similar plateaux in brazil. the lower surface formed during the neogene after the deposition of the sabiá formation and an uplift event that raised the higher surface to its present elevation around 1000 m a.s.l. the uplift resulted in rejuvenation of the relief and subsequent formation of the lower surface. progressive backward erosion along the main rivers has resulted in escarpments that separate the two surfaces. the escarpments are pronounced at geological boundaries with large differences of erosional resistance. even the lower surface is presently under destruction due to minor subsequent uplift. in summary, we find that the passive margin topography in the study area was shaped many millions of years after the early cretaceous break-up of the south atlantic. the conclusion that the landscape is mainly cenozoic is thus in agreement with that of, e.g. king (1967). a better understanding of the timing of uplift events will be achieved from apatite fission-track data as well as the amount of exhumation involved in the formation of the erosion surfaces. acknowledgements the study was funded by statoilhydro do brasil and petrobras. references bonow, j.m., lidmar-bergström, k. & japsen, p. 2006: palaeosurfaces in central west greenland as reference for identification of tectonic movements and estimation of erosion. global and planetary change 50, 161–183. bonow, j.m., japsen, p., green, p.f., wilson, r.w., chalmers, j.a., klint, k.e.s., van gool, j.a.m., lidmar-bergström, k. & pedersen, a.k. 2007: a multi-disciplinary study of phanerozoic landscape development in west greenland. geological survey of denmark and greenland bulletin 13, 33–36. cobbold, p.r., meisling, k.e. & mount, v.s. 2001: reactivation of an obliquely rifted margin, campos and santos basins, southeastern brazil. american association of petroleum geologists bulletin 85, 1925–1944. cobbold, p.r., marais-gilchrist, g., chiossi, d., fonseca chaves, f., gomes de souza, f. & lilletveit, r. 2008: large submarine slides on a steep and narrow continental margin (camamu basin, ne brazil). american association of petroleum geologists annual convention, san antonio, texas, 20–23 april, 1 p. cprm-geological survey of brazil 2001: geological map of brazil, 1:5 000 000 (cd-rom). cprm-geological survey of brazil 2003: geologia e recursos minerais do estado da bahia 1:1 000 000 (cd-rom). davison, i. 1999: tectonics and hydrocarbon distribution along the brazilian south atlantic margin. geological society, london, special publications 153, 133–151. gallagher, k., hawkesworth, c.j. & mantovani, m.s.m. 1994: the denu dation history of the onshore continental margin of se brazil inferred from apatite fission track data. journal of geophysical research 99, 18117–18145. japsen, p., bonow, j.m., green, p.f., chalmers, j.a. & lidmar-bergström, k. 2006. elevated, passive continental margins: long-term highs or neogene uplifts? new evidence from west greenland. earth and planetary science letters 248, 315–324. japsen, p., bonow, j.m., green, p.f., chalmers, j.a., & lidmar-bergström, k. 2009: formation, uplift and dissection of planation surfaces at passive continental margins. earth surface processes and landforms 34, 683–699. jarvis, a., reuter, h. i., nelson, a. & guevara, e. 2008: hole-filled seamless srtm data v4, international centre for tropical agriculture (ciat). available from http://srtm.csi.cgiar.org. king, l.c. 1967: the morphology of the earth, 2nd edition, 699 pp. edinburgh: oliver and boyd. lima, m. da guia 2008: a história do intemperismo na província bor borema oriental, nordeste do brasil: implicações paleoclimáticas e tectônicas, 251 pp. unpublished ph.d. thesis, universidade federal do rio grande do norte, brazil. magnavita, l.p., davison, i. & kusznir, n.j. 1994: rifting, erosion, and uplift history of the recôncavo-tucano-jatobá rift, northeast brazil. tectonics 13, 367–388. morais neto, j.m., green, p.f., karner, g.d. & alkmim, f.f. in press: age of the serra do martins formation, borborema plateau, northeastern brazil: constraints from apatite and zircon fission track analysis. boletim de geociências da petrobras 16. ollier, c.d. 1985: morphotectonics of passive continental margins: introduction. zeitschrift für geomorphologie n.f., suppl. 54, 1–9. sant’anna, l.g., schorscher, h.d. & riccomini, c. 1997: cenozoic tectonics of the fonseca basin region, eastern quadrilátero ferrífero, mg, brazil. journal of south american earth sciences 10, 275–284. spier, c.a., vasconcelos, p.m. & oliviera, s.m.b. 2006: 40ar/39ar geo chronological constraints on the evolution of lateritic iron deposits in the quadrilátero ferrífero, minas gerais, brazil. chemical geology 234, 79–104. viana, c.f., gama junior, e., simões, i.a., moura, j.a., fonseca, j.r. & alves, r.j. 1971: revisão estratigráfica da bacia recôncavo/tucano. boletim técnico da petrobras 14, 157–192. 84 authors’ addresses j.m.b & p.j., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: jbon@geus.dk p.f.g., geotrack international pty ltd, 37 melville road, brunswick west 3055, victoria, australia. p.r.c., géosciences-rennes (umr6118 du cnrs), université de rennes, 35042 rennes cedex, france. a.j.p., geological survey of brazil (cprm), avenida ulysses guimaraes, 2862, 41213-000 salvador, brazil. r.l., statoilhydro, angola team, stavanger, norway. d.c., statoilhydro do brazil, rio de janeiro, brazil. rosa_2008:rosa-2008 01/07/09 15:49 side 84 data article | short karlsson et al. 2023: geus bulletin 53. 8338. https://doi.org/10.34194/geusb.v53.8338 1 of 8 a data set of monthly freshwater fluxes from the greenland ice sheet’s marine-terminating glaciers on a glacier–basin scale 2010–2020 nanna b. karlsson*1 , kenneth d. mankoff1,2,3 , anne m. solgaard1 , signe h. larsen1 , penelope r. how1 , robert s. fausto1 , louise s. sørensen4 1geological survey of denmark and greenland (geus), copenhagen, denmark. 2autonomic integra llc, new york, ny, usa. 3nasa goddard institute for space studies, new york, ny, usa. 4dtu space – national space institute, technical university of denmark, kgs. lyngby, denmark abstract the loss of mass from the greenland ice sheet causes an increasing influx of freshwater to the greenlandic fjords and the oceans. freshwater fluxes from marine-terminating glaciers are important to understand fjord circulation and ecosystem dynamics. here, we present a data set constructed by reformulating existing products into a shared temporal and spatial framework. we combine three publicly available data sets of solid-ice discharge (iceberg), liquid-surface runoff (runoff) and basal melt to present a cohesive overview of the flow of freshwater from marine-terminating glaciers to the greenlandic fjords. we also calculate glacier drainage basins and compare our findings to previous studies showing that drainage-basin sizes may vary considerably depending on how they were reconstructed. the data set will be a valuable asset to oceanographic, glaciological and marine biological research activities. data the greenland ice sheet discharges significant volumes of freshwater into the fjords and oceans (mankoff et al. 2020). this freshwater discharge is known to modify and influence the physical, chemical and biological properties of the fjords and coastal seas (e.g. hopwood et al. 2020). the combined volume of freshwater that exits from marine-terminating glaciers is currently not readily available on a glacier–basin scale due to disparate data sets. this study presents a data set constructed by combining three publicly available *correspondence: nbk@geus.dk received: 01 dec 2022 revised: 21 apr 2023 accepted: 12 jun 2023 published: 22 aug 2023 keywords: ice mass loss, ice sheet – fjord interactions, ice surface melt, iceberg discharge, basal melt abbreviations dem: digital elevation model gimp: greenland ice mapping project k2021: karlsson et al. (2021) mar: modèle atmosphérique régional measures: making earth system data records for use in research environments m2019: mankoff et al. (2019) m2020: mankoff et al. (2020) promice: programme for monitoring of the greenland ice sheet racmo: regional atmospheric climate model rcm: regional climate model smb: surface mass balance geus bulletin (eissn: 2597–2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: marit-solveig siedenkrantz (university of aarhus, denmark) reviewed by: taryn black (university of washington, usa) and two anonymous reviewers funding: see page 7 competing interests: see page 7 additional files: see page 7 tabular abstract geographical coverage greenland, ice-sheet margins temporal coverage 2010–2020 (inclusive); monthly resolution subject(s) cryosphere, oceanography, atmosphere and climate data format(s) modelled and reformatted data as csv files are available here: https://doi.org/10.22008/fk2/bovbvr sample collection & analysis the three terms of the freshwater flux are obtained in the following way: ice discharge: from remotely sensed observations of ice velocity and thickness. the latter is based on a combination of remote sensing and models. surface melt runoff: from regional climate models. basal melt: from a combination of remote-sensing data and mathematical models. parameters ice discharge (icebergs), surface melt and basal melt related publications mankoff et al. 2019, 2020; karlsson et al. 2021 potential application(s) for these data to quantify the freshwater flux for each glacier; compare different sources of freshwater; resolve the seasonal variability in freshwater flux for different fjords. https://doi.org/10.34194/geusb.v53.8338 https://orcid.org/0000-0003-0423-8705 https://orcid.org/0000-0001-5453-2019 https://orcid.org/0000-0002-8693-620x https://orcid.org/0000-0002-3656-3521 https://orcid.org/0000-0002-8088-8497 https://orcid.org/0000-0003-1317-8185 https://orcid.org/0000-0002-3771-4061 mailto:nbk@geus.dk https://doi.org/10.22008/fk2/bovbvr karlsson et al. 2023: geus bulletin 53. 8338. https://doi.org/10.34194/geusb.v53.8338 2 of 8 www.geusbul let in.org data sets of solid-ice discharge (iceberg), liquid-surface runoff (runoff) and basal melt from the greenland ice sheet. this new product merges disparate data sets into a product that shares spatial and temporal resolution, enabling easy comparison of mass-loss processes on glacier–basin scales. the following sections give a brief overview of the data sets that form the foundation of the product presented here. solid-ice discharge we use the term ‘solid-ice discharge’ to describe the ice mass that is lost at the marine margin as either icebergs, bergy bits or submarine melt. while this term is traditionally reported as a mass flux in kg or gt, we convert it to liquid water equivalents to be consistent with the other volume-loss terms. the solid-ice discharge from more than 200 flux gates situated near the front of marine-terminating glaciers was compiled and presented by mankoff et al. (2019). here, we give a brief overview of their methods and results and refer to the original manuscript for details. mankoff et al. (2019; m2019 for the remainder of this manuscript) calculate ice discharge by considering the mass flow rate through predefined flux gates. the method uses ice thickness and an estimated vertical velocity distribution based on the observed surface velocity, vs, to calculate the discharge. this may be formulated as: dg = ρvhw (1) where dg is the ice discharge across a gate, ρ is the average density of ice, v is the depth-averaged horizontal velocity perpendicular to the gate, h is the ice thickness and w is the gate width. in m2019, the velocity v(z) is assumed to be equal at all ice depths, implying that observed surface velocities, vs, represent depth-averaged velocities (see enderlin et al. 2014; king et al. 2018), thus assuming v = vs. in the gate-discharge method, the gate should ideally be at the grounding line of the outlet glacier, where the ice loses contact with its bed and begins to float. in m2019, however, gates are located 5 km from the calving front (which is co-located with the grounding line on most greenlandic outlet glaciers). the choice of a 5 km distance from the calving front is to keep surface melt between the gate and the front small, while also far enough upstream to minimise errors in bed-location ice thickness that tend to increase towards the calving front (see m2019). for outlet glaciers with ice shelves (as identified in morlighem et al. 2017), the flux gate is located 5 km from the grounding line. the flux gates are generated by an algorithm ensuring that gate locations can be reproduced or adapted if a glacier outlet changes flow speed or configuration. ice-flow velocity data are based on data generated from sentinel 1a and 1b derived by promice (solgaard et al. 2021), measures 0478 and measures 0646 (mouginot et al. 2018a, b; howat 2017a). the bed topography that provides ice thicknesses at the gates is from bedmachine v3 (morlighem et al. 2017). the solid-ice discharge data set spans from 1986 to the present day with a temporal resolution ranging from sub-annual to weekly. this data set has a spatial resolution of 200 m at the discharge gates (mankoff et al. 2019). surface-meltwater runoff the surface-meltwater runoff is the liquid water that drains from the ice sheet as a result of melt and rainfall. the term runoff implies that water retained on the ice sheet is not part of the estimate. the runoff stems from model outputs from regional climate models (rcms). more information on methods can be found in mankoff et al. (2020), in the following referred to as m2020. again, we give a brief overview of the methods and results here and refer to the original mankoff et al. (2020) manuscript for details. m2020 estimated greenland’s liquid-water discharge from surface melt by routing rcm runoff estimates from all points on the ice sheet to the ice margin and coastal outlets. the routing was derived from an ice-sheet surface digital elevation model (dem), an ice-sheet bed dem, an ice-sheet mask (citterio & ahlstrøm 2013), a land-surface dem and an ocean mask from the greenland ice mapping project (gimp): land ice and ocean classification mask, version 1 (howat et al. 2014; howat 2017b). m2020 used arcticdem v3 100m (porter et al. 2018) as the surface dem while ice thickness from bedmachine v3 (morlighem et al. 2017) was used to calculate the subglacial routing. the m2020 data set contains two runoff estimates using two rcm outputs based on the modèle atmosphérique régional (mar; fettweis et al. 2017) and the regional atmospheric climate model (racmo; noël et al. 2018), respectively. in both cases, the runoff, r, is defined by: r = me + ra – rt − rf (2) where me is melt, ra is rainfall, rt is retention and rf is refreezing. the rcm outputs were re-gridded to the same 1 km grid using offline statistical downscaling (noël et al. 2016; fettweis et al. 2020). the model outputs were validated against 10 river-discharge time series with a daily resolution. the final data product from m2020 was daily liquid discharge values from 24 507 ice marginal outlets (that is, ice runoff that discharges at the margin, either on land at the land – water boundary, or subglacially at the ice – ocean boundary) and 29 635 land coast outlets (that is, land-sourced runoff that discharges either at the coast or subglacially). in this work, we only use the ice-sourced runoff that discharges at the ice – ocean boundary. https://doi.org/10.34194/geusb.v53.8338 http://www.geusbulletin.org karlsson et al. 2023: geus bulletin 53. 8338. https://doi.org/10.34194/geusb.v53.8338 3 of 8 www.geusbul let in.org basal meltwater the meltwater specifically caused by subglacial melting, i.e. melt that is separate from surface meltwater, has been quantified by karlsson et al. (2021), hereafter k2021. again, we refer the reader to the original manuscript for more details. in brief, k2021 combined remote-sensing products and modelling to constrain three sources of basal heat in order to obtain local basal-melt rates, bm: bm = e /(ρl) (3) where e is available energy for the melt, ρ is ice density and l is the latent heat of fusion. the first heat source, the geothermal flux, was constructed as the mean of three different heat-flux models. the second heat source, the frictional heat, was obtained from the full-stokes ice-flow model elmer/ice (gillet-chaulet et al. 2012) that emulates present-day ice dynamics by minimising the misfit between modelled and observed surface velocities. the last heat source, viscous-heat dissipation, was estimated by converting the gravitational potential energy of surface meltwater into heat, assuming that all water reaches the bed, essentially providing an upper limit on the available energy. in k2021, surface-meltwater volumes were obtained from the ensemble mean of the greenland smb intercomparison project (fettweis et al. 2020) and routed to the bed using bedmachine v3. the first and second heat sources were masked by an independent estimate of where the ice sheet was likely to be thawed or frozen based on a combination of models and data (macgregor et al. 2016). in k2021, the basal-meltwater data set was published as maps specifying basal-melt rates in m yr–1 for 1 km grid cells and estimates of total basal-meltwater volumes were given mainly as regional totals. in contrast to the two preceding data products, the basal meltwater is only in part temporally changing. the geothermal flux and the friction heat are assumed to be constant while the viscous heat varies with meltwater input. changes to the original data set for this study in m2020, the runoff was routed to the bed of the ice using the topography from bedmachine v3. here, the topography is updated to bedmachine v4. in this study, we only present runoff derived from mar and we refer readers to m2020 for a detailed discussion of the similarities between mar and racmo. in k2021, the viscous-heat dissipation was given as the melt resulting from the annual average surface-meltwater volumes for 1995–2010. in this updated data product, the basal melt due to viscous heat has been calculated based on the temporally varying surface-meltwater volumes. thus, the geothermal and friction terms are constant in time, while the viscous-heat dissipation is resolved monthly using the surface-meltwater volumes from mar. statistical analysis and data processing we combine the three data sets using the solid-ice discharge gates as our starting point (fig. 1, yellow lines). the discharge gates are associated with individual glacier outlets or (in some cases) individual tributaries of glaciers. we then tie the surface-meltwater runoff to the discharge gate by identifying individual streamlines that intersect discharge gates (fig. 1, blue lines). the outlets of these streamlines (fig. 1, blue dots) are then assigned as belonging to the discharge gate, and the n72°30′ 72°0′ 71°30′ –52°0′ –50°0′ –48°0′ 0 15 30 45 km solid-ice discharge gates surface-water outlets surface-water routing surface-water basins basel melt rates (m/yr) ≤ 0.2 0.2 – 0.4 0.4 – 0.6 0.6 – 0.8 0.8 – 1.0 1.0 – 1.2 1.2 – 1.4 1.4 – 1.6 1.6 – 1.8 1.8 – 2.0 salliarutsip sermia umiammakku sermiat kangilliup sermia fig. 1 example from west greenland showing the three data products. yellow lines indicate the locations of solid-ice discharge gates from m2020. the surface-meltwater product is shown as blue streamlines with water outlets as purple dots and their associated basins in black. the basal-melt product is shown in red as m/yr. background image from satellite background image is from sentinel-2 (european space agency), 20 august 2022. https://doi.org/10.34194/geusb.v53.8338 http://www.geusbulletin.org karlsson et al. 2023: geus bulletin 53. 8338. https://doi.org/10.34194/geusb.v53.8338 4 of 8 www.geusbul let in.org corresponding surface-water basin (fig. 1, black lines) is associated with the discharge gate. the outline of the surface-water basins was published as part of the m2020 data set. discharge gates frequently have more than one surface-water basin associated with them; for example, wide glaciers often have several streamlines intersecting their discharge gates. note that the data presented here only include the surface meltwater that exits via a marine-terminating glacier. freshwater sources such as land-terminating glaciers, rivers and lakes are not included in our estimate. for these fluxes, we refer to m2020. for the basal-melt component, we calculate the route of each individual grid cell in the basal-melt map to the margin of the ice sheet. when a route intersects a discharge gate, the grid cell and all the grid cells intersected along the route, are assigned as belonging to the discharge gate in question. the subglacial routes are assumed to follow the steepest gradient in the hydropotential (shreve 1972): φ = ρw gzb + ρi g (zs + zb) (4) where ρw is the density of water, ρi is the density of ice, and zb and zs are the elevations of bed and surface topography, respectively. this methodology closely resembles the one employed by m2020 when calculating the surface water basins with some exceptions. here, the surface and bed topographies stem from bedmachine v3 and are smoothed by several ice thicknesses (10 km by 10 km). we perform this smoothing based on several considerations, primarily that the bed topography from bedmachine (regardless of version) is highly uncertain in some areas, which can cause erroneous hydropotential lows leading to the apparent formation of subglacial lakes. this is in contrast with observations showing that there is a limited number of subglacial lakes under the greenland ice sheet (livingstone et al. 2022), implying that most subglacial water flows to the margin of the ice sheet in agreement with theoretical considerations of surface and bed slopes (pattyn 2008). another consideration is the fact that subglacial water may flow in sediments under the ice; thus, hydropotential lows may be circumvented by water if the bedrock material is porous. we suggest that it is highly likely that water will find a way to travel from areas of high pressure to areas of low pressure (ice-sheet interior to margin). thus, topographic smoothing is necessary to force water routes to exit at the margin. in contrast to the basins for the basal melt, the surface-melt basins are not calculated based on a smoothed topography. this difference allows us to explore the impact of topographic smoothing on our results and we discuss this further in the section uncertainties. we resample d and r to the same timescale, a monthly volume loss. we chose this timescale as a pragmatic common time. r is model-derived and available daily. the temporal resolution of d varies from seasonal to sub-monthly depending on satellite acquisition timing. where necessary, d is linearly interpolated in time if the data are not available monthly. data description and main features our data set represents 267 individual discharge gates. the freshwater flux from each discharge gate is accessible as a csv file with the naming convention glaciername_region_gate.csv, where the glacier name is the official greenlandic name (following bjørk et al. 2015 and the oqaasileriffik placename database maintained by asiaq), the region follows the naming convention from mouginot et al. (2019c; see fig. 3) and ‘gate’ is the gate id from m2019. the data set is accompanied by a metadata text file and a file containing geographical information and naming convention for each gate. here, we present the contents of the data set using the three glaciers shown in fig. 1 as an example. figure  2a, b, c show the monthly surface runoff, solid-ice discharge, and basal melt for the glaciers salliarutsip sermia, umiammakku sermiat and kangilliup sermia (rink isbræ) for 2010–2020. all three glaciers experienced large surface melt in 2012 and again in 2019. in comparison, the solid-ice discharge has lower variability – although kangilliup sermia has distinct variations in solid-ice discharge during the year, often with a winter minimum. note that the y-axis is different by an order of magnitude for kangilliup sermia. this is clearly seen in fig. 2d. the temporal resolution of our data makes it possible to consider the seasonality of the volume loss. figure 2e shows the summer and winter volume loss, where summer is defined as the months april through september while winter is defined as october through march. most volume loss happens during the summer months for all three glaciers, but kangilliup sermia has a larger proportion of winter volume loss, including liquid volume loss, compared to its neighbours. according to our full data set, all marine-terminating glaciers in greenland lose more volume during the summer than the winter. however, the distribution of the volume loss varies. for example, 32 glaciers have a winter volume loss that is close to summer volume loss, including glaciers arfalluarfiup sermia (diebitsch, nw region; regions indicated in fig. 3), apuseeq qinnilik (ce region) and nunatakassaap sermia (alison gletscher, nw region). at the other end of the scale, 16 glaciers experience 95% of their volume loss during the summer. this includes glaciers apuseeq (køge  bugt, https://doi.org/10.34194/geusb.v53.8338 http://www.geusbulletin.org karlsson et al. 2023: geus bulletin 53. 8338. https://doi.org/10.34194/geusb.v53.8338 5 of 8 www.geusbul let in.org regions n and se), waltershausen gletscher (ne region) and kangiata nunaata sermia (sw region). we can expand on this analysis and consider regional differences. figure 3 shows the seasonal distribution of volume loss from all discharge gates reported from the seven regions of greenland. all regions experience more volume loss during the summer than during the winter but with some regional differences. the regions se and cw experience 38% of their volume loss during the winter months while in comparison winter volume loss from sw and no regions accounts for 20% and 26% of annual volume loss, respectively. the regions ne and sw have the largest proportion of winter volume loss as liquid, 6% and 7%, respectively. table 1 details the average monthly volume loss from each region during the winter and summer months. table 2 shows regional annual averages of solidand liquid-volume loss. uncertainties the uncertainties associated with our data set fall into two categories: (1) those inherited from the data sets upon which we have constructed ours and (2) those associated with constructing the drainage basins. for the former, we adopt the uncertainties as stated in the m2019, m2020 and k2021 data sets. following m2019, the ice-discharge uncertainties are within 10%. uncertainties primarily stem from unknown or poorly sampled ice thicknesses and uncertainties in the surface-velocity data. we refer the reader to m2019, appendix a for a thorough discussion of their treatment of uncertainties. the runoff is influenced by several different kinds of uncertainties, including uncertainty in the delay between melt (at a location on the ice sheet) and discharge of the water at the margin, uncertainties associated with the rcms that generate the meltwater volume, and the uncertainty in the basin delineation. m2020 discusses these uncertainties in detail but refrains from stating a global uncertainty, noting that depending on local glacier settings, the uncertainties likely vary substantially between sites. particularly small basins are likely to be subject to substantial errors due to basin delineation. the stated uncertainty for the rcms is 15%. we advise users of our data set to consider this as a lower boundary. the uncertainty associated with the basal-melt data set is also spatially variable and stems from the fact that basal conditions of the greenland ice sheet are widely unknown, including the poorly constrained geothermal flux. k2021 uses an asymmetrical uncertainty range to capture the full possibility of basal conditions. here, for simplicity and to be consistent with m2019 and m2020, average volume loss (m3/yr) 1e10 (d) surface melt ice discharge basal melt 0 20 40 60 80 100 % volume loss salliarutsip sermia umiammakku sermiat kangilliup sermia (e) volume loss distribution summer solid summer liquid winter solid winter liquid jan-2010 jan-2015 jan-2020 0 1 2 3 4 5 6 m 3 1e8 (a) salliarutsip sermia surface melt ice discharge basal melt jan-2010 jan-2015 jan-2020 0 1 2 3 4 5 6 m 3 1e8 (b) umiammakku sermiat jan-2010 jan-2015 jan-2020 date 0.00 0.25 0.50 0.75 1.00 1.25 1.50 m 3 1e9 (c) kangilliup sermia average mass loss (gt/yr) 0.00 0.25 0.50 0.75 1.00 1.25 1.50 0.0 2.0 4.0 8.0 10.0 12.0 14.06.0 fig. 2 monthly volume loss terms as attributed to different processes for glaciers. (a) salliarutsip sermia. (b) umiammakku sermiat. (c) kangilliup sermia (locations in fig. 1). we partition the volume loss into surface melt (dark blue), iceberg calving (light blue) and basal melt (magenta). for comparison are (d) average annual volume loss for 2010–2020 and (e) relative seasonal distribution of volume loss for the same period. official greenlandic names from bjørk et al. (2015). https://doi.org/10.34194/geusb.v53.8338 http://www.geusbulletin.org karlsson et al. 2023: geus bulletin 53. 8338. https://doi.org/10.34194/geusb.v53.8338 6 of 8 www.geusbul let in.org we use the largest uncertainty as a conservative estimate at 21%. again, we refer readers to k2021 for a more detailed discussion of the different uncertainties. combining the uncertainty ranges listed above using root-sum square, we estimate the overall uncertainty to be + +10% 15% 21%2 2 2 = 28% as a lower boundary, although we note that uncertainties are not independent, even though we treat them as such. the uncertainty associated with constructing the drainage basins is also thoroughly discussed in m2020 in connection with the surface-meltwater product. to further investigate the uncertainty, we compare the basins constructed in this work based on smoothed surface and bed topography to the basins from m2020. by far the biggest difference between drainage basins is found in north-east greenland. in m2020, the outlet named ce cw ne no nw se sw s w (a) 0 20 40 60 80 100 % volume loss ce cw ne no nw se sw (b)summer liquid summer solid winter liquid winter solid fig. 3 seasonality of all regions in greenland. (a) summer (brown) and winter (purple) volume loss for each region where the size of the circles indicates the total volume loss. the background image shows ice-flow velocities from measures (joughin 2020). (b) same as (a) but as bars. the proportion of the volume loss that is liquid is shown with darker colours and a black outline. the seven regions are defined by mouginot et al. 2019c as follows: south-west (sw), central west (cw), north-west (nw), north (no), north-east (ne), central east (ce), and south-east (se). table 1 volume loss in solid and liquid form region average monthly solid flux (october–march), 107 m3 average monthly liquid flux (october–march), 107 m3 average monthly solid flux (april–september), 107 m3 average monthly liquid flux (april–september), 107 m3 ce 12 ± 3 0.4 ± 0.1 12 ± 3 13 ± 4 cw 45 ± 12 1.8 ± 0.5 46 ± 13 33 ± 9 ne 29 ± 8 1.9 ± 0.5 29 ± 8 33 ± 9 no 11 ± 3 0.4 ± 0.1 12 ± 3 22 ± 6 nw 13 ± 4 0.3 ± 0.1 13 ± 4 11 ± 3 se 13 ± 4 0.4 ± 0.1 13 ± 4 10 ± 3 sw 12 ± 3 0.9 ± 0.3 12 ± 3 37 ± 10 total 136 ± 38 6.1 ± 1.7 138 ± 38 159 ± 45 the table shows the average flux during 2010–2020 for the winter months (october through march) and summer months (april through september) in units of 107 m3. the seven regions are defined in fig. 3: south-west (sw), central west (cw), north-west (nw), north (no), north-east (ne), central east (ce), and south-east (se). https://doi.org/10.34194/geusb.v53.8338 http://www.geusbulletin.org karlsson et al. 2023: geus bulletin 53. 8338. https://doi.org/10.34194/geusb.v53.8338 7 of 8 www.geusbul let in.org zachariae isstrøm drains a large part of the region, while in this study the main outlet is spaltegletsjer ( nioghalvfjerdsfjorden). the change is so large that the basin size changes by 98% for zachariæ. the second largest difference in basin size occurs for kjer gletsjer in north-west greenland, which shrinks in size by 6.7e10 m2 or several multitudes of its original size. supplementary table 1 summarises the 10 largest differences in basins between m2020 and this study. although the values look prohibitively large, we note that 28% of basins change in size by less than 20%. due to the large variability in the basins discussed here, we do not quantify the uncertainty related to the basins, noting that small basins are likely to be particularly uncertain. acknowledgments nbk thanks n. damgaard for helpful insights and advice on figure designs. the authors thank t. black for a very thorough and constructive review that greatly improved this manuscript. we also thank two anonymous reviewers for valuable comments and suggestions, and handling editor m-s seidenkrantz. funding statement this work was supported by the polar+ 4d greenland project (2020–2022), which was funded by the european space agency (esa) via esa contract no. 4000132139\i-ef and by promice, which is funded by the geological survey of denmark and greenland (geus) and the danish ministry of climate, energy and utilities under the danish cooperation for environment in the arctic (dancea), and is conducted in collaboration with dtu space (technical university of denmark) and asiaq, greenland. competing interests we declare no competing interests. author contributions nbk: conceptualisation, data curation, formal analysis, funding acquisition, methodology, visualisation, writing. kdm: data curation, methodology, validation, writing. ams: data curation, methodology, funding acquisition, methodology, writing. shl: methodology, validation, writing. prh: resources, software, writing. rsf: funding acquisition, project administration, writing. lss: funding acquisition, project administration, writing. additional files supplementary table 1 along with modelled and reformatted data are available at https://doi.org/10.22008/fk2/bovbvr references bjørk, a.a., kruse, l.m. & michaelsen, p.b. 2015: brief communication: getting greenland’s glaciers right – a new data set of all official greenlandic glacier names. the cryosphere 9(6), 2215–2218. https://doi. org/10.5194/tc-9-2215-2015 citterio, m. & ahlstrøm, a.p. 2013: brief communication ‘the aerophotogrammetric map of greenland ice masses’. the cryosphere 7(2), 445–449. 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b., scheuchl, b. & wood, m. 2019c: forty-six years of greenland ice sheet mass balance from 1972 to 2018. proceedings of the national academy of sciences 116, 9239–9244. https://doi.org/10.1073/pnas.1904242116 noël, b., van de berg, w.j., machguth, h., lhermitte, s., howat, i., fettweis, x. & van den broeke, m.r. 2016: a daily, 1 km resolution data  set of downscaled greenland ice sheet surface mass balance (1958–2015). the cryosphere 10(5), 2361–2377. https://doi. org/10.5194/tc-10-2361-2016 noël, b. et al. 2018: modelling the climate and surface mass balance of polar ice sheets using racmo2 – part 1: greenland (1958–2016). the cryosphere 12, 811–831. https://doi.org/10.5194/tc-12-811-2018 pattyn, f. 2008: investigating the stability of subglacial lakes with a full stokes ice-sheet model. journal of glaciology 54(185), 353–361. https://doi.org/10.3189/002214308784886171 porter, c. et al. 2018: arcticdem v3. havard dataverse. https://doi. org/10.7910/dvn/ohhukh shreve, r.l. 1972: movement of water in glaciers. journal of glaciology 11(62), 205–214. https://doi.org/10.3189/s002214300002219x solgaard, a. et al. 2021: greenland ice velocity maps from the promice project. earth system science data 13(7), 3491–3512. https://doi. org/10.5194/essd-13-3491-2021 https://doi.org/10.34194/geusb.v53.8338 http://www.geusbulletin.org https://doi.org/10.5194/essd-12-2811-2020 https://doi.org/10.5194/essd-12-2811-2020 https://doi.org/10.1002/2017gl074954 https://doi.org/10.7280/d1mm37 https://doi.org/10.7280/d1gw91 https://doi.org/10.1073/pnas.1904242116 https://doi.org/10.5194/tc-10-2361-2016 https://doi.org/10.5194/tc-10-2361-2016 https://doi.org/10.5194/tc-12-811-2018 https://doi.org/10.3189/002214308784886171 https://doi.org/10.7910/dvn/ohhukh https://doi.org/10.7910/dvn/ohhukh https://doi.org/10.3189/s002214300002219x https://doi.org/10.5194/essd-13-3491-2021 https://doi.org/10.5194/essd-13-3491-2021 a data set of monthly freshwater fluxes from the greenland ice sheet’s marine-terminating glaciers o data solid-ice discharge surface-meltwater runoff basal meltwater changes to the original data set for this study statistical analysis and data processing data description and main features uncertainties acknowledgments funding statement competing interests author contributions additional files references figures fig. 1 example from west greenland showing the three data products. yellow lines indicate the locat fig. 2 monthly volume loss terms as attributed to different processes for glaciers. (a) salliarutsip fig. 3 seasonality of all regions in greenland. (a) summer (brown) and winter (purple) volume loss f tables table 1 volume loss in solid and liquid form. table 2 annually averaged volume loss in solid and liquid form (2010-2020). geological survey of denmark and greenland bulletin 11, 87-99 87 zircon geochronology from the kangaatsiaq– qasigiannguit region, the northern part of the 1.9–1.8 ga nagssugtoqidian orogen, west greenland kristine thrane and james n. connelly the kangaatsiaq–qasigiannguit region in the northern part of the palaeoproterozoic nagssugtoqidian orogen of west greenland consists of poly-deformed orthogneisses and minor occurrences of interleaved, discontinuous supracrustal belts. laser ablation icp-ms 207pb/206pb analyses of detrital zircons from four metasedimentary rocks (supplemented by ion probe analysis of one sample) and igneous zircons from six granitoid rocks cutting metasedimentary units indicate that the supracrustal rocks in the kangaatsiaq–qasigiannguit (christianshåb) region are predominantly archaean in age. four occurrences of metasedimentary rocks are clearly archaean, two have equivocal ages, and only one metasedimentary unit, from within the naternaq (lersletten) supracrustal belt, is demonstrably palaeoproterozoic and readily defines a large fold complex of this age at naternaq. the 2.9–2.8 ga ages of detrital archaean grains are compatible with derivation from the local basement orthogneisses within the nagssugtoqidian orogen. the detrital age patterns are similar to those of metasediments within the central nagssugtoqidian orogen but distinct from age patterns in metasediments of the rinkian belt to the north, where there is an additional component of pre-2.9 ga zircons. synkinematic intrusive granitoid rocks constrain the ages of some archaean deformation at 2748 ± 19 ma and some palaeoproterozoic deformation at 1837 ± 12 ma. keywords: nagssugtoqidian orogen, deformation, la-icp-ms, zircon, metasediment ______________________________________________________________________________________________________________________________________________________________________________________________________ k.t., geological institute, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: kthrane@geol.ku.dk j.n.c., department of geological science, university of texas at austin, austin, tx 78712, usa. the kangaatsiaq–qasigiannguit area that is the focus of this paper (fig. 1) forms a large part of the northern nagssugtoqidian orogen, which is interpreted as the southern part of a major collisional orogenic system that crops out in central and northern west greenland and adjacent parts of eastern canada (connelly et al. 2006). the northern nagssugtoqidian orogen, which was re-investigated in 2001–2003 by the geological survey of denmark and greenland (geus) in co-operation with external partners, is underlain by poly-deformed, variably reworked grey archaean orthogneiss interleaved with dismembered archaean and palaeoproterozoic supracrustal rocks of volcanic and sedimentary origin. only few significant time marker horizons (such as distinct suites of mafic dykes or one or more groups of characteristic supracrustal rocks with known ages) are present. the primary objectives of this geochronological study were therefore to determine the extent of palaeoproterozoic metasedimentary rocks and attempt to directly date different phases of deformation and metamorphism. a number of lithological, structural and metamorphic features, especially in the kangaatsiaq–aasiaat area (fig. 1), © geus, 2006. geological survey of denmark and greenland bulletin 11, 87–99. available at: www.geus.dk/publications/bull 88 provide some immediate constraints on the archaean and palaeoproterozoic geological evolution of the northern nagssugtoqidian orogen, and are outlined here as an introduction to the geochronological study. the study area may be divided into two different tracts based on metamorphism and structural style (piazolo et al. 2004; mazur et al. 2006, this volume). the tract south-west and southeast of kangaatsiaq is metamorphosed at granulite facies grade and is characterised by a general wsw–ene-trending structural grain with large, moderately to steeply plunging fold structures and undeformed to weakly deformed, synkinematic granitic neosome (van gool et al. 2002; garde 2004). several e–w-trending mafic dykes occur south of kangaatsiaq around 68°n. they are undeformed and discordant to the main structures and lithological boundaries, but variably metamorphosed (glassley & sørensen 1980; árting 2004). these dykes are presumed to be of palaeoproterozoic age and perhaps related to pre-nagssugtoqidian rifting (árting 2004), and if so would constrain the deformation and granulite facies metamorphism south of kangaatsiaq to be archaean in age, whereas the thermal event recorded by the dykes themselves would be palaeoproterozoic. the remainder of the study area, to the north and east of kangaatsiaq, is at amphibolite grade (e.g. hollis et al. 2006, this volume), and does not display any signs of retrogression from granulite facies except within a c. 10 km thick transition zone adjacent to the granulite facies terrain. these northern and eastern areas generally possess a much more intense planar and linear tectonic fabric than in the south, commonly including a strong subhorizontal extension lineation that also penetrates the late granitic neosome. furthermore, a structural discordance occurs in the naternaq area (fig. 1) between wnw–ese-trending amphibolite to the west and the structurally overlying, ne–sw-trending naternaq supracrustal belt in the east, suggesting that the respective structures of the two supracrustal units are unrelated to each other and of different age (mazur et al. 2006, this volume). in addition, the northern and eastern areas also host occasional mafic dykes on islands north-east of aasiaat and on the southern coast of sydostbugten (fig. 1). although these dykes are still largely coherent and unmigmatised, they are intensely deformed, almost concordant with their host rocks, and tectonically thinned to about 1–2 m thick. both the granitic neosome, the naternaq supracrustal rocks, and the deformed dykes provide relative age constraints on the intense deformation in the northern and eastern parts of the study area. if it is again assumed that the deformed dykes in the north are palaeoproterozoic, it would follow 69° 54° 68° inland ice palaeogene basalt quaternary deposits metasedimentary rocks palaeoproterozoic (naternaq supracrustal belt) arfersiorfik quartz diorite amphibolite archaean variably reworked. may include palaeoproterozoic supracrustal components granodioritic and granitic gneiss orthogneiss metasedimentary rocks amphibolite qasigiannguit aasiaat kangaatsiaqkangaatsiaq ka ng ers un eq saqqarputataneq naternaq/ lersletten sydostbugten 20 km amitsoq amitsoq amitsoq kangaatsiaq 5151°°51° 480041 483631 448392 448394 480054 470515 463129 463257 448004 464435 500 km fig. 1. simplified geological map of the kangaatsiaq–qasigiannguit region, with sample locations. 89 448004 metasediment ne of kangaatsiaq 1 22419 0.21046 2909 7.7 2 30428 0.20444 2862 4.9 3 27946 0.20011 2827 5.9 4 41739 0.20405 2859 4.8 5 66864 0.20548 2870 5.8 6 52963 0.20846 2894 4.2 7 70422 0.18216 2673 3.9 8 46147 0.20058 2831 5.2 9 73807 0.20449 2862 3.7 10 41046 0.20919 2899 5.9 11 75396 0.20183 2841 6.5 12 84403 0.20242 2846 4.2 13 27932 0.20390 2858 7.5 14 102678 0.20107 2835 6.0 15 129026 0.18972 2740 6.9 16 46526 0.20716 2883 5.5 17 71802 0.20579 2873 6.3 18 659551 0.20033 2829 4.9 19 165828 0.19242 2763 6.1 20 73874 0.19895 2818 5.4 21 20371 0.18508 2699 5.5 22 21982 0.17386 2595 5.3 23 33164 0.21021 2907 6.4 24 86486 0.18406 2690 5.1 25 99148 0.20434 2861 4.4 26 54554 0.20049 2830 5.6 27 52435 0.17947 2648 3.8 28 45325 0.20018 2828 4.3 29 75391 0.18983 2741 3.0 30 77047 0.20775 2888 4.7 31 32181 0.19856 2814 3.1 32 50218 0.19811 2811 3.3 33 28491 0.20148 2838 4.3 34 21801 0.19768 2807 4.6 35 22466 0.20144 2838 5.9 36 48990 0.20013 2827 3.3 37 26535 0.20445 2862 5.8 38 114112 0.18364 2686 2.1 39 41973 0.20083 2833 3.4 40 57549 0.20210 2843 2.6 41 26869 0.25312 3204 2.7 42 49986 0.19811 2811 3.0 43 20567 0.19924 2820 4.6 44 5467 0.19417 2778 9.2 45 29042 0.20106 2835 4.2 46 9255 0.18565 2704 6.7 47 18597 0.20277 2849 4.3 48 112700 0.17900 2644 2.3 49 84858 0.19285 2767 2.8 50 14867 0.17617 2617 8.2 51 24908 0.19765 2807 3.7 52 11178 0.19419 2778 5.5 53 91255 0.17807 2635 4.3 54 28338 0.20395 2858 4.7 55 24855 0.18937 2737 4.8 56 7761 0.18757 2721 9.0 57 12436 0.19563 2790 6.5 58 12571 0.19849 2814 5.4 448392 granite, kangersuneq 1 318831 0.19069 2748 3.5 2 361175 0.18986 2741 2.7 3 406129 0.19945 2822 3.3 4 163666 0.18800 2725 3.6 5 205935 0.19380 2775 3.1 6 227282 0.19148 2755 3.0 7 123658 0.18830 2727 3.1 8 294942 0.18738 2719 5.1 9 407109 0.19415 2778 5.0 10 481882 0.19342 2771 5.1 11 212302 0.18706 2716 4.7 12 407464 0.19766 2807 3.0 13 491512 0.19950 2822 4.2 14 94032 0.18905 2734 3.0 44 19762 0.19659 2798 6.5 45 36375 0.11197 1832 9.0 46 17548 0.18024 2655 11.7 47 9064 0.19962 2823 10.9 48 11530 0.19199 2759 7.8 49 64234 0.11111 1818 10.3 50 7566 0.16671 2525 11.0 51 38371 0.10619 1735 9.1 52 15464 0.18496 2698 9.1 53 15534 0.14855 2329 9.3 54 27618 0.11461 1874 8.4 55 15853 0.19890 2817 9.4 56 10902 0.18316 2682 11.0 57 8905 0.20952 2902 9.4 58 6493 0.18680 2714 10.5 59 16248 0.19267 2765 8.7 60 43736 0.14480 2285 6.4 61 48133 0.11266 1843 12.3 62 52656 0.12037 1962 10.7 63 88558 0.11566 1890 7.2 64 65717 0.10988 1797 8.8 65 14964 0.15276 2377 8.5 66 59777 0.11635 1901 7.3 67 23849 0.16781 2536 8.6 68 16825 0.19548 2789 12.3 69 17260 0.14321 2266 8.8 70 28538 0.19178 2757 10.3 71 65355 0.11192 1831 8.4 72 58939 0.12654 2050 9.6 73 69509 0.11151 1824 6.8 74 49893 0.13319 2140 8.6 75 76541 0.11448 1872 8.1 76 67611 0.11351 1856 8.0 77 58817 0.10691 1747 8.9 78 69641 0.11468 1875 11.3 79 57943 0.11116 1818 4.7 463129 synkinematic granite, saqqarput 1 2283739 0.18759 2721 3.0 2 1266859 0.19890 2817 3.6 3 344375 0.18904 2734 4.1 4 531257 0.18883 2732 5.5 5 565174 0.19441 2780 4.6 6 247679 0.18844 2729 4.9 7 460352 0.20656 2879 3.8 8 1058356 0.19510 2786 3.2 9 405853 0.18296 2680 6.8 10 554081 0.17660 2621 3.0 11 1363965 0.20876 2896 4.0 12 935603 0.19609 2794 2.7 13 605520 0.18625 2709 2.6 14 516471 0.17880 2642 5.1 15 822166 0.18954 2738 5.5 16 655511 0.20436 2861 4.0 17 755390 0.19587 2792 5.0 18 127005 0.18594 2707 5.9 19 422927 0.19261 2765 4.8 20 653950 0.19183 2758 4.0 21 1135912 0.19679 2800 2.9 22 1096015 0.19497 2785 2.4 23 653314 0.19170 2757 2.6 24 228993 0.20892 2897 7.4 25 221510 0.21878 2972 9.2 26 346159 0.20045 2830 6.6 27 37048 0.16944 2552 11.9 28 97269 0.17974 2650 10.8 29 85210 0.18521 2700 8.7 30 172975 0.18213 2672 8.2 31 171017 0.17446 2601 9.0 32 165639 0.18002 2653 5.5 33 133709 0.18851 2729 7.8 34 182742 0.18315 2682 5.2 35 125199 0.18194 2671 7.0 36 169034 0.19244 2763 7.5 37 85182 0.16778 2536 9.4 15 173984 0.18392 2689 4.9 16 109933 0.18320 2682 5.7 17 50994 0.19030 2745 4.6 18 13465 0.18864 2730 6.2 19 21101 0.17697 2625 6.4 20 99478 0.18431 2692 5.8 21 108982 0.18687 2715 3.9 22 33255 0.18969 2739 5.7 23 192965 0.18985 2741 6.8 24 280376 0.18859 2730 6.1 25 287725 0.18761 2721 5.1 26 101087 0.18975 2740 4.7 27 85412 0.18695 2716 4.3 28 125421 0.18639 2711 6.2 29 45496 0.18854 2730 5.7 30 75470 0.18108 2663 4.9 31 40021 0.18483 2697 4.9 32 48564 0.19087 2750 5.6 33 101631 0.18754 2721 3.2 34 117730 0.18292 2679 2.5 35 49211 0.18807 2725 3.8 36 112930 0.18847 2729 2.9 37 202637 0.18826 2727 2.2 38 102831 0.18893 2733 2.2 39 55266 0.19174 2757 2.8 40 45349 0.18639 2711 4.5 41 104416 0.18742 2720 3.3 42 188642 0.18481 2696 2.2 43 97849 0.18797 2724 2.6 44 145946 0.17859 2640 4.0 448394 metasediment, kangersuneq 1 695411 0.16006 2456 3.6 2 52863 0.20417 2860 5.3 3 443404 0.11180 1829 3.8 4 1104248 0.11717 1913 4.4 5 579803 0.11532 1885 4.7 6 231966 0.12351 2008 4.9 7 362937 0.10979 1796 3.1 8 303451 0.11334 1854 3.9 9 277637 0.11090 1814 3.1 10 380962 0.12251 1993 3.9 11 115636 0.16115 2468 6.0 12 461330 0.11598 1895 3.5 13 493717 0.19532 2787 3.7 14 207687 0.17408 2597 5.3 15 393460 0.11112 1818 2.7 16 242631 0.10929 1788 4.9 17 132013 0.19838 2813 4.9 18 178574 0.10859 1776 6.1 19 303683 0.10774 1762 4.1 20 241139 0.15569 2409 7.4 21 51484 0.12627 2047 5.5 22 53211 0.12080 1968 7.7 23 53531 0.11660 1905 6.9 24 23896 0.19593 2793 6.2 25 79706 0.11458 1873 6.4 26 89308 0.11277 1845 6.3 27 63147 0.11176 1828 8.6 28 12586 0.19431 2779 6.5 29 100961 0.13800 2202 7.0 30 12592 0.16853 2543 9.7 31 57066 0.11124 1820 8.5 32 51764 0.11438 1870 10.1 33 218430 0.11295 1847 7.4 34 78169 0.10868 1777 9.7 35 96773 0.11345 1855 8.4 36 143102 0.11037 1806 6.7 37 33237 0.13986 2225 6.4 38 92186 0.10819 1769 9.7 39 80332 0.12802 2071 8.3 40 35510 0.19386 2775 8.0 41 11684 0.18283 2679 7.4 42 21389 0.19620 2795 8.2 43 14702 0.19528 2787 9.7 table 1. zircon la-icp-ms 207pb-206pb data spot 206pb (cps) 207pb/206pb age (ma) 2σ %spot 206pb (cps) 207pb/206pb age (ma) 2σ % spot 206pb (cps) 207pb/206pb age (ma) 2σ % 90 table 1 (continued) spot 206pb (cps) 207pb/206pb age (ma) 2σ %spot 206pb (cps) 207pb/206pb age (ma) 2σ % spot 206pb (cps) 207pb/206pb age (ma) 2σ % 38 101924 0.18113 2663 7.6 39 178145 0.20233 2845 8.9 40 35771 0.18557 2703 8.3 41 23866 0.19033 2745 10.0 42 78620 0.22740 3034 11.3 43 88660 0.20740 2885 7.8 44 264622 0.18784 2723 6.5 45 219347 0.19029 2745 5.2 46 215885 0.18648 2711 9.4 47 117234 0.17897 2643 9.0 48 29425 0.18849 2729 8.9 49 59458 0.18317 2682 10.0 50 44706 0.19343 2772 11.5 51 47511 0.20920 2899 14.7 52 52220 0.18199 2671 10.6 53 58597 0.17855 2639 12.7 54 47047 0.18403 2690 6.8 55 85100 0.19408 2777 8.0 56 150838 0.18512 2699 9.1 57 42220 0.18152 2667 8.4 58 182300 0.17444 2601 6.3 59 181857 0.19158 2756 5.8 60 656182 0.18446 2693 3.8 61 1625878 0.19433 2779 4.0 62 1033362 0.19004 2742 7.1 463257 metasediment, amitsoq 1 94060 0.20316 2852 4 2 126332 0.19611 2795 4 3 245874 0.21538 2946 4 4 154787 0.20172 2840 4 5 68216 0.19982 2825 4 6 428099 0.18536 2701 4 7 100464 0.18737 2719 4 8 106081 0.19619 2794 4 9 178863 0.20711 2883 4 10 72018 0.19715 2808 4 11 102336 0.18203 2671 4 12 218440 0.21010 2907 4 13 144435 0.20135 2837 4 14 73200 0.20617 2875 4 15 103837 0.17190 2575 4 16 100257 0.15132 2361 4 17 120574 0.20468 2864 4 18 88636 0.20786 2889 4 19 129707 0.20434 2861 4 20 133092 0.21288 2928 4 21 176443 0.20885 2897 4 22 48385 0.19377 2774 5.2 23 61486 0.19210 2760 8.2 24 60275 0.22028 2983 3.6 25 44550 0.20247 2846 6.8 26 30952 0.20191 2842 5.9 27 7392 0.18113 2663 8.0 28 20571 0.19351 2772 8.6 29 15357 0.19909 2819 6.7 30 31747 0.20691 2881 7.0 31 22596 0.20038 2829 8.1 32 35613 0.20326 2853 7.7 33 15204 0.20568 2872 10.5 34 24804 0.19424 2778 8.2 35 17027 0.19728 2804 5.8 36 21866 0.19494 2784 8.4 37 13148 0.20815 2891 9.3 38 27502 0.20821 2892 6.1 39 26648 0.20452 2863 8.3 40 27263 0.18701 2716 10.7 41 46499 0.14869 2331 6.7 42 36471 0.20245 2846 6.7 43 64587 0.20102 2834 8.7 44 31986 0.20332 2853 8.9 45 46174 0.21292 2928 8.9 46 61284 0.20775 2888 6.6 47 41436 0.20439 2862 6.8 48 72510 0.20430 2861 5.1 49 67124 0.20347 2854 7.9 50 37221 0.19866 2815 7.0 14 162431 0.11214 1834 4.2 15 42424 0.11157 1825 5.5 16 109374 0.11367 1859 4.3 17 52552 0.11218 1835 6.4 18 67018 0.11372 1860 5.2 19 38574 0.11173 1828 5.3 20 184531 0.11162 1826 2.5 21 420598 0.11279 1845 2.5 22 675700 0.11206 1833 1.7 480041 quartz diorite, qasigiannguit 1 130744 0.19864 2815 5.2 2 127784 0.19639 2796 4.7 3 30286 0.19684 2800 4.6 4 28769 0.20120 2836 4.7 5 41484 0.19738 2805 4.7 6 19837 0.20222 2844 7.0 7 35270 0.19516 2786 5.5 8 48190 0.20117 2836 5.6 9 35231 0.19155 2756 5.0 10 93023 0.19849 2814 4.8 11 36116 0.19465 2782 4.8 12 82370 0.19490 2784 10.4 13 58322 0.19959 2823 7.2 14 61328 0.19932 2821 6.7 15 58943 0.19951 2822 11.6 16 25024 0.19642 2797 3.7 17 61436 0.19260 2765 5.7 18 56976 0.19778 2808 7.2 19 43379 0.19673 2799 5.3 20 68466 0.19216 2761 7.7 480054 tonalite intruding mafic complex 1 8174 0.20380 2857 10.3 2 10419 0.19353 2772 7.4 3 18592 0.20474 2864 5.5 4 213185 0.15919 2447 5.9 5 17238 0.19681 2800 5.3 6 11204 0.20444 2862 6.4 7 17603 0.20266 2848 7.2 8 36568 0.20791 2889 9.5 9 42107 0.20580 2873 6.5 10 30492 0.19888 2817 7.1 11 25591 0.20251 2847 7.2 12 14106 0.20374 2856 6.4 13 18218 0.20025 2828 4.4 14 20117 0.19869 2816 5.8 15 42704 0.20836 2893 6.8 16 29583 0.20336 2853 7.6 17 23900 0.19738 2805 9.0 483631 granite intruding metasediment 1 259347 0.19195 2759 3.1 2 155094 0.19256 2764 2.4 3 94380 0.19393 2776 3.3 4 33706 0.18539 2702 3.7 5 57040 0.18930 2736 3.2 6 42007 0.19204 2760 3.5 7 71104 0.19592 2793 3.8 8 33803 0.19616 2794 4.1 9 79316 0.19110 2752 3.9 10 13550 0.18486 2697 7.3 11 57313 0.19407 2777 4.3 12 56128 0.19155 2756 3.1 13 56620 0.19591 2792 3.4 14 101430 0.19250 2764 3.6 15 41779 0.19211 2760 3.4 16 367814 0.19593 2793 2.0 17 118779 0.18771 2722 2.7 18 26473 0.16345 2492 13.9 19 45218 0.19626 2795 3.6 20 76399 0.19586 2792 3.6 21 24409 0.18865 2730 4.7 464435 metasediment, naternaq 1 620630 0.12031 1961 4 2 184109 0.14293 2261 4 3 177844 0.12013 1959 4 4 508546 0.12356 2019 4 5 521502 0.11208 1833 4 6 560514 0.11545 1887 4 7 455565 0.10976 1795 4 8 294092 0.11916 1944 4 9 407244 0.12342 2006 4 10 217026 0.11556 1889 4 11 240050 0.11948 1948 4 12 285058 0.12258 1995 4 13 651523 0.11831 1938 4 14 367681 0.12552 2036 4 15 337595 0.12582 2041 4 16 465635 0.11732 1917 4 17 321964 0.12267 1995 4 18 1150971 0.11774 1922 4 19 266990 0.12383 2013 4 20 374761 0.12502 2030 4 21 297903 0.12321 2002 4 22 87395 0.11806 1927 4.3 23 88706 0.12335 2005 6.5 24 53359 0.13734 2194 6.5 25 72472 0.11727 1915 5.8 26 16174 0.13058 2106 4.4 27 125924 0.11860 1935 4.6 28 81398 0.11940 1947 6.3 29 93208 0.11402 1864 6.3 30 131504 0.12001 1956 4.2 31 107212 0.12410 2016 6.6 32 66177 0.12687 2055 7.7 33 72338 0.11608 1897 4.1 34 163168 0.11450 1872 4.4 35 92298 0.12172 1982 5.5 36 220614 0.12305 2001 4.3 37 112247 0.12249 1993 5.7 38 76097 0.11662 1905 4.7 39 76847 0.11907 1942 4.1 40 39338 0.11918 1944 6.8 41 51130 0.11856 1935 6.6 42 86970 0.12690 2055 5.0 43 128237 0.11865 1936 5.3 44 47126 0.12236 1991 5.5 45 84328 0.11894 1940 5.6 46 307611 0.12127 1975 5.0 47 115975 0.11797 1926 6.9 48 149629 0.12005 1957 7.5 49 385671 0.11759 1920 7.7 50 454534 0.12076 1967 6.8 51 245715 0.11875 1937 7.4 52 320479 0.12160 1980 6.4 53 1098422 0.10860 1776 7.8 54 578352 0.12047 1963 4.1 55 186864 0.11804 1927 7.0 56 79299 0.11614 1898 5.5 57 255968 0.10900 1783 3.3 58 707747 0.11876 1938 7.4 59 97872 0.12173 1982 6.5 60 109798 0.11649 1903 5.4 61 489175 0.12003 1957 4.0 470515 pegmatite ne of kangaatsiaq 1 149998 0.11231 1837 3.3 2 114249 0.11111 1818 2.3 3 128681 0.11066 1810 2.5 4 249115 0.11405 1865 2.5 5 174468 0.11307 1849 2.2 6 188664 0.11253 1841 2.7 7 118304 0.11223 1836 3.0 8 69708 0.11414 1866 4.3 9 107108 0.11379 1861 3.5 10 57749 0.11019 1803 5.2 11 255076 0.11091 1814 2.7 12 112190 0.11350 1856 4.7 13 145495 0.11459 1873 3.2 91 that the northern and eastern parts of the study area are strongly reworked by nagssugtoqidian deformation and amphibolite facies metamorphism. geochronological targets and methods both archaean and palaeoproterozoic supracrustal sequences are known to exist within the nagssugtoqidian orogen (marker et al. 1999; nutman et al. 1999). depositional ages of such supracrustal belts may be constrained by the ages of detrital zircons in their sedimentary components, since the youngest grains define their maximum age of deposition; conversely, the age of a magmatic rock that has intruded the supracrustal sequence may serve to define a lower depositional age limit. ideally, both metasediment and cross-cutting magmatic rocks from the same outcrop should be analysed to best constrain the timing of deposition. however, cross-cutting intrusive rocks of appropriate age (i.e. other than late palaeoproterozoic pegmatites) were not generally present. the direct dating of archaean or palaeoproterozoic deformation by dating e.g. synkinematic granitoids requires the rather scarce occurrence of an intrusive rock that unequivocally both cuts and is affected by a single fabric. ten samples from the kangaatsiaq–qasigiannguit area, mainly provided by members of the geus mapping groups in 2001–2002, have been analysed by quadrupole laser ablation inductively coupled mass spectrometry (la-icp-ms) at the university of texas at austin (zircon pb-pb data, table 1); additional ion probe u-pb zircon data from one metasedimentary rock were obtained at the nordsim laboratory, naturhistoriska riksmuseet, stockholm (table 2). analytical details are given in the appendix. the samples were collected from seven different, dismembered metasedimentary sequences (four samples of metasediment and four samples of cross-cutting granitoid rocks), and from intrusive granite and orthogneiss that constrain the timing of deformation (two samples). all ages of rocks presented in this manuscript have been calculated using isoplot/ex (ludwig 1999) and are reported with 2-sigma uncertainties. the main advantage of using the la-icp-ms technique for zircon geochronology is that each analysis only lasts about two minutes, whereas the analytical time on an ion microprobe is typically around 15–20 minutes. this becomes an important factor when analysing detrital rocks, where analysis of a large number of detrital grains is essential to achieve good statistics. the major limitation of the la-icp-ms technique employed is that u-pb ratios could not be measured, and only 207pb/206pb ages are obtained. a direct indication of concordance therefore is not available, and all the ages obtained should be interpreted conservatively to represent minimum ages of crystallisation or metamorphism. furthermore, common pb corrections cannot be carried out due to interference in the plasma of 204hg from the carrier gases. a test of the la-icp-ms instrument used in this study was carried out by connelly et al. (2006), who analysed zircons from the itilli diorite, disko bugt, west greenland using both la-icpms and id-tims methods and found that the 207pb/206pb age of 3019 ± 23 ma obtained with the laser instrument compared well with its id-tims age of 3030 +8/–5 ma. 1 404 161 167 0.40 0.37 0.1160 0.36 5.390 1.81 0.3370 1.77 –1.4 2 384 170 170 0.44 0.02 0.1162 0.33 5.727 1.80 0.3575 1.77 4.4 3 235 121 101 0.51 1.34 0.1159 0.75 5.338 1.93 0.3341 1.78 –2.2 4 411 178 173 0.43 0.24 0.1170 0.33 5.483 1.81 0.3398 1.77 –1.5 5 489 195 208 0.40 0.24 0.1158 0.32 5.534 1.80 0.3466 1.77 1.6 6 360 123 149 0.34 0.02 0.1174 0.33 5.511 1.80 0.3406 1.77 –1.6 7 416 163 176 0.39 0.04 0.1160 0.31 5.537 1.80 0.3463 1.77 1.3 8 413 178 178 0.43 0.03 0.1169 0.31 5.629 1.80 0.3494 1.77 1.4 9 653 281 239 0.43 0.70 0.1137 0.40 4.743 1.82 0.3025 1.77 –9.5 10 237 88 99 0.37 0.04 0.1163 0.39 5.500 1.81 0.3429 1.77 0.0 11 239 108 101 0.45 0.05 0.1174 0.39 5.521 1.81 0.3410 1.77 –1.6 12 448 224 194 0.50 0.03 0.1170 0.39 5.539 1.81 0.3434 1.77 –0.5 13 2192 1713 975 0.78 0.15 0.1165 0.17 5.337 1.78 0.3323 1.77 –3.3 1898 6 1935 16 1970 30 1894 13 1875 17 1858 29 1911 6 1898 16 1886 29 1892 6 1906 16 1918 29 1916 6 1902 16 1890 29 1895 6 1906 16 1917 29 1909 6 1921 16 1932 30 1860 7 1775 15 1704 27 1901 7 1901 16 1900 29 1896 7 1883 16 1872 29 1918 7 1904 16 1891 29 1911 7 1907 16 1903 29 1903 3 1875 15 1849 2 table 2. zircon ion probe u-th-pb data from sample 464435, naternaq spot u th pb th/u f 206 % 207pb σ % 207pb σ % 206pb σ % disc. % # ppm ppm ppm measured 206pb 235u 238u (conv.) 207pb σ 207pb σ 206pb σ 206pb 235u 238u ages (ma) errors on ratios and ages are quoted at 1σ level. f 206 %: the fraction of common 206pb, estimated from the measured 204pb. disc. % (conv.): degree of discordance of the zircon analysis (at the centre of the error ellipse). 92 fig. 2. zircon age data from the kangaatsiaq–qasigiannguit region. a, c, d, g, j, k: weighted average plots of la-icp-ms 207pb/206pb age data of igneous rocks. b, e, h, i: probability density plots of la-icp-ms 207pb/206pb age data of metasediments. f: ion probe u-pb age data (concordia plot), sample 464435. intercepts at 1904 ± 8 and 452 ± 290 ma mswd = 1.9 1740 1780 1820 1860 1900 1980 2020 0.29 0.31 0.33 0.35 0.37 464435 metasediment, naternaq n = 13 464435 metasediment, naternaq n = 61 0 2 4 6 8 10 12 14 16 fr eq ue nc y fe 207pb/206pb age (ma) 20 6 p b/ 23 8 u 207pb/235u 4.64.2 5.0 5.4 5.8 6.21700 2100 2500 2900 3300 0 2 4 6 8 10 12 1700 2100 2500 2900 3300 fr eq ue nc y 448394 metasediment, kangersuneq n = 79 b 207pb/206pb age (ma) 2400 2600 2800 3000 3200 mean age = 2801 ± 34 ma mswd = 0.094 a 20 7 p b/ 20 6 p b ag e (m a) 480041 quartz diorite, qasigiannguit n = 20 2600 2800 3000 3200 480054 tonalite intruding mafic complex n = 16 mean age = 2839 ± 46 ma mswd = 0.117 2500 2600 2700 2800 2900 3000 mean age = 2723 ± 15 ma mswd = 0.33 dc 20 7 p b/ 20 6 p b ag e (m a) 20 7 p b/ 20 6 p b ag e (m a) 448392 granite, kangersuneq n = 41 93 geochronological age constraints of metasedimentary belts the study area contains numerous, dismembered, discontinuous belts of metasedimentary rocks that may be either archaean or palaeoproterozoic in age. while the main focus of this work was to constrain their timing of deposition, provenance information gained through the detrital zircons permits regional correlation of these metasedimentary belts. the analysed samples are presented in tables 1–2 and fig. 2 and discussed below from north to south. sample 480041, quartz diorite intruding metasedimentary and metavolcanic rocks at qasigiannguit sample 480041 of a homogeneous, grey, medium-grained quartz diorite was collected 3 km east of qasigiannguit at 68°48.83′n, 51°08.05′w (fig. 1). the rock consists of plagioclase, quartz, hornblende and biotite and has a strong linear fabric. the quartz diorite forms a 3–4 km long elongate body exposed on the top of the ridge facing qasigiannguit. its contact relationships are generally equivocal due to deformation, but at the south-western margin the contact appears to be intrusive into a metasedimentary-metavolcanic sequence. the zircons from this sample are clear and stubby. twenty grains were analysed, which yield consistent 207pb/ 206pb ratios corresponding to a weighted mean age of 2801 ± 34 ma (mswd = 0.094, fig. 2a). the age is interpreted as the crystallisation age of the quartz diorite, implying that the supracrustal sequence it cuts must also be archaean. samples 448394, metasediment and 448392, intruding granite on the south coast of kangersuneq sample 448394 (68°46.24′n, 50°52.16′w) from a pelitic metasedimentary rock and sample 448392 (68°46.20′n, 50°51.55′w) of a granite that cuts the metasedimentary belt, were collected c. 200 m apart on the south coast of kangersuneq (fig. 1). after the analytical results were obtained, the locality was revisited in 2003 and the previously reported field relations confirmed (jeroen van gool, personal communication 2003). the zircons from the metasediment are brownish, elongate, 100–200 µm long, and in many cases cracked and showing clear signs of dissolution. the least altered and, presumably, least disturbed zircons were chosen for anal1700 2100 2500 2900 3300 fr eq ue nc y 463257 metasediment, amitsoq n = 50 i 207pb/206pb age (ma) 0 2 4 6 8 10 12 14 483631 granite intruding metasediment n = 21 2200 2400 2600 2800 3000 mean age = 2763 ± 20 ma mswd = 0.47 g 20 7 p b/ 20 6 p b ag e (m a) 463129 synkinematic granite, saqqarput n = 62 20 7 p b/ 20 6 p b ag e (m a) mean age = 2748 ± 19 ma mswd = 1.14 j 2300 2500 2700 2900 3100 3300 207pb/206pb age (ma) 1700 2100 2500 2900 3300 fr eq ue nc y 448004 metasediment ne of kangaatsiaq n = 58 h 0 2 4 6 8 10 12 14 470515 pegmatite ne of kangaatsiaq n = 22 mean age = 1837 ± 12 ma mswd = 0.av44 k 20 7 p b/ 20 6 p b ag e (m a) 1720 1760 1800 1840 1880 1920 1960 94 yses. in several cases it was possible to distinguish broad rims containing more u than the cores, and in such cases both core and rim were analysed. seventy-nine spots were analysed and yield an age spectrum with a large peak at 1850 ma and a smaller one at 2800 ma (fig. 2b). a range of intermediate ages (2500–2100 ma) between the two peaks are also present (see below). the archaean peak comprises only analyses from cores, whereas the 1850 ma peak consists of analyses from both cores and rims. the zircons from the granite are typically brown, and larger than those in the metasediment. in size they range from 100–350 µm and occur both as slender and somewhat stubby crystals. core–rim zonation is observed in the majority of the zircons. of 44 grains analysed from the granite, 41 yielded a consistent pattern of 207pb/206pb ratios corresponding to an average age of 2723 ± 15 ma (mswd = 0.33) (fig. 2c); both cores and rims were analysed in several grains without observing any age variations. the remaining three grains yield ages from 2822 to 2807 ma and are most likely inherited. due to the homogeneity of the zircon population it is highly unlikely that the zircons are detrital grains inherited from the metasediment. they are also unlikely to have been inherited from the orthogneisses adjacent to the metasediment, as these do not generally contain such young zircons. therefore, the age of 2723 ± 15 ma is interpreted as the emplacement age of the granite, and the metasediment must also be of archaean age. the 1850 ma peak for the zircons in the metasediment is therefore interpreted to date nagssugtoqidian metamorphism, and the 2500–2100 ma ages most likely represent archaean zircons that have suffered pb-loss; alternatively the latter analyses might represent accidental mixtures of cores and rims. sample 480054, tonalite intruding mafic complex c. 25 km south-east of kangersuneq sample 480054, a biotite-hornblende tonalite, was collected from a relatively undeformed tonalitic body c. 2 km2 in size at 68°35.53′n, 50°30.28′w within a large mafic supracrustal complex near the inland ice about 25 km south-east of kangersuneq (fig. 1). the tonalite is light grey in colour, mediumto coarse-grained, homogeneous, and has a weak linear fabric. it is feldspar-phyric with phenocrysts up to 2 cm in diameter. near the contacts with the surrounding mafic rocks the tonalite contains xenoliths of the mafic supracrustal rocks, and its intrusive nature is unequivocal. dykes of tonalite, 50 cm to 2 m wide, extend from the main tonalite body into rocks of the surrounding large mafic complex. the tonalite sample yielded a population of large, euhedral, mostly prismatic, clear to yellow zircons. seventeen spots on zircon grains were analysed, and sixteen of them generated a consistent spectrum of 207pb/206pb ratios corresponding to an average age of 2839 ± 46 ma (mswd = 0.117) (fig. 2d). the consistent 207pb/206pb ratios indicate that little or no pb-loss has occurred. the age result is therefore interpreted to closely reflect the crystallisation age of the tonalite, and the mafic complex intruded by the tonalite must consequently also be archaean. sample 464435, metasedimentary rock from the naternaq supracrustal belt naternaq (lersletten) is an extensive quaternary outwash plain with scattered outcrops of precambrian basement gneisses and supracrustal rocks (østergaard et al. 2002). a sample of very fine-grained mica schist was collected from the extensive naternaq supracrustal belt at 68°24.10′n, 51°56.70′w (fig. 1). the zircons are elongate, 50–150 µm long, and vary in colour from clear to slightly brown. all 61 analyses carried out yield palaeoproterozoic 207pb/206pb ages ranging from 2261 to 1776 ma, with the main peak of 207pb/206pb ages around 1950 ma (fig. 2e). while it is possible that some of the younger grains may be metamorphic, we interpret the majority of the grains to be detrital because of the igneous appearance of the zircons and because they are older than any metamorphic event so far described in the nagssugtoqidian orogen (e.g. connelly et al. 2000, earliest metamorphism at c. 1870 ma). furthermore, it seems unlikely that a zircon population from a metasedimentary rock would only comprise metamorphic grains and not contain a single detrital grain. consequently this requires a palaeoproterozoic deposition age for the metasedimentary unit at naternaq. in order to confirm the obtained 207pb/206pb la-icpms ages, zircons from this sample were also analysed on the cameca ims 1270 ion microprobe at the nordsim laboratory, swedish museum of natural history, stockholm. the thirteen ion probe analyses yield a cluster of ages on the concordia diagram of fig. 2f (table 2), with an upper intercept isochron age of 1904 ± 8 ma (mswd = 1.9). the ion probe age is thus slightly younger than the c. 1950 ma peak obtained by la-icp-ms, but the two data sets overlap within the large analytical error of the latter method, and the apparent older laicp-ms age could be due to common pb for which we were unable to correct. in conclusion, the ion probe data clearly demonstrate that the zircons are older than any metamorphic ages hitherto obtained from the nagssug95 toqidian orogen. the most likely interpretation is that the sediment at naternaq was derived from the magmatic arc that formed in the central part of the orogen between 1920 and 1870 ma (kalsbeek et al. 1987; kalsbeek & nutman 1996; whitehouse et al. 1998; connelly et al. 2000). the new age data from the naternaq supracrustal belt have important consequences for the structural interpretation of the naternaq area, documenting the existence of large palaeoproterozoic folds. sample 483631, granite vein intruding metasediment on strike with the naternaq supracrustal sequence a sample (483631) of an 8 cm wide vein of muscovitebiotite granite vein was collected east of naternaq at 68°31.04′n, 51°17.87′w (fig. 1). the granite vein is deformed but clearly intrudes mica schist on strike with the eastern part of the naternaq supracrustal belt. most of the zircons in sample 483631 are long, prismatic grains with a distinct core-rim zonation and range from clear to brownish in colour. twenty-two analyses were carried out, mostly on cores, but also on a few rims. all except one analysis yielded consistent 207pb/206pb ratios corresponding to an average age of 2763 ± 20 ma (mswd = 0.47, fig. 2g). one core analysis yielded a 207pb/206pb age of 2467 ± 158 ma, which may be due to pb-loss. the analyses demonstrate that the granite is archaean. the regional basement does not generally contain orthogneisses with such young ages and, therefore, limits the possibility that the zircons in the granite were inherited. it follows that the supracrustal rocks east of naternaq on strike with the naternaq supracrustal belt must also be of archaean age. sample 448004, metasedimentary belt near kangaatsiaq sample 448004 of a quartz-rich metasedimentary rock was collected north-east of kangaatsiaq at 68°21.15′n, 53°13.18′w (fig. 1). the zircons vary from elongate to stubby but all have been rounded during transport. they are clear and between 50 to 200 µm in length. a weak igneous zonation is present in the majority of the grains. fifty-eight spots were analysed, and both core and rim were analysed in several grains. the age population ranges from 2909 to 2595 ma, with a single grain yielding an age of 3200 ma (fig. 2h). a peak is present around 2850 ma, which is a common age for polyphase orthogneisses within the nagssugtoqidian orogen (kalsbeek et al. 1987; kalsbeek & nutman 1996; whitehouse et al. 1998; connelly & mengel 2000; hollis et al. 2006, this volume). the scatter of younger ages from 2800 to 2595 ma most likely results from variable degrees of pb loss. we cannot with certainty assign an age of deposition to this sediment, as the data are compatible with both archaean and palaeoproterozoic sedimentation. sample 463257, quartz-rich metasedimentary sequence at amitsoq sample 463257 of a mediumto fine-grained, quartz-rich metasedimentary rock with abundant small garnets was collected at the head of the fjord amitsoq at 68°05.78′n, 52°30.99′w (fig. 1). it is part of an extensive metasedimentary sequence, which is significantly more quartz-rich than all other metasedimentary rocks reported from the kangaatsiaq map area, although similar rocks have been observed in the central nagssugtoqidian orogen (jeroen van gool, personal communication 2003). the zircons are clear, 50–250 µm in length and vary from elongate to stubby, and have been rounded during transport. most zircons show clear igneous zonation, and some contain a high-u metamorphic rim. this rim was unfortunately too thin to analyse with the la-icp-ms. all except two of the 50 analysed grains yield archaean ages, with two exceptions interpreted to have suffered pbloss. the zircons show very little age variation, with a major peak of 207pb/206pb ages at around 2850 ma (fig. 2i). this age is comparable to that of the surrounding regional orthogneisses, and indicates that the detritus may be locally derived. however, it is only possible to conclude that the sediment was deposited after 2850 ma. dating of deformation using synkinematic granitic rocks sample 463129, synkinematic granite vein at saqqarput a sample of synkinematic granite (463129) was collected at saqqarput at 68°09.22′n, 52°42.65′w (fig. 1). the granite occurs as fineto medium-grained, subconcordant veins in the orthogneiss. sixty-two analyses yielded 207pb/ 206pb ratios corresponding to an average 207pb/206pb age of 2748 ± 19 ma (mswd = 1.14; fig. 2j). this is interpreted to be the crystallisation age of the granite since, as in sample 483631 described above, it is unlikely that zircons 96 of this age are inherited. some of the observed deformation in the host basement orthogneisses must, therefore, also be archaean in age. sample 470515, pegmatite north-east of kangaatsiaq a sample of pegmatite (470515) was collected north-east of kangaatsiaq at 68°20.23′n, 59°00.17′w (fig. 1). the pegmatite forms a 020° trending vertical sheet cutting orthogneiss. the pegmatite is a member of a conjugate set of pegmatites within the kangaatsiaq map area that indicate late, regional n–s orientated compression (ian alsop, personal communication 2002), and itself contains evidence of ductile sinistral shear along its margins; the regional foliation is deflected into sinistral shear fabrics, indicating that the pegmatite emplacement took place later than the foliation formation in the gneisses, but while the host rock was still hot enough to behave in a ductile manner. the sample contains large, brownish, prismatic zircons that vary from slender to short and stubby in shape. twenty-two spots were analysed on thirteen grains. both rims and cores were analysed on several grains, but no age variation was documented between the two. all analyses yield the same result within uncertainties; the average 207pb/ 206pb ratio corresponds to an age of 1837 ± 12 ma (mswd = 0.44) (fig. 2k). this is interpreted to be the emplacement age of the pegmatite, and is considered to date the late nagssugtoqidian n–s compression. discussion and conclusions the zircon ages obtained from this study show that metasedimentary rocks in the kangaatsiaq–qasigiannguit region are predominately archaean in age. the best age constraints come from the four archaean granitoid rocks that cut four different supracrustal belts. only one metasediment was analysed from these belts and yields an archaean detrital peak of c. 2800 ma and a palaeoproterozoic peak of c. 1850 ma. the palaeoproterozoic peak is attributed to the growth of metamorphic zircon during nagssugtoqidian metamorphism. similarly, the occurrences of ages between 2800–1850 ma are attributed to pb-loss from detrital grains and/or mixed analysis of detrital and metamorphic zircon. two of the three remaining metasedimentary rocks yield only archaean detrital ages with peaks between 2800 and 2900 ma, which are similar to the ages of the basement orthogneisses in the nagssugtoqidian orogen (kalsbeek et al. 1987; kalsbeek & nutman 1996; whitehouse et al. 1998; connelly & mengel 2000) and may indicate that the sedimentary sequences were derived from local sources. while it is tempting to conclude that these metasediments might also themselves be of archaean age, the lack of cross-cutting granites from these locations only permits the conclusion that they must be younger than 2850 ma. thus, the sediments could have been deposited either in the archaean at around 2850–2750 ma (i.e., before the regional 2.75 ga metamorphism documented within the nagssugtoqidian orogen), or during the palaeoproterozoic (most likely before the formation of the arfersiorfik–sisimiut arc; kalsbeek et al. 1987; kalsbeek & nutman 1996; whitehouse et al. 1998; connelly et al. 2000). the remaining sample of metasediment (464435), collected from the naternaq supracrustal belt, is the only metasediment in this study which is interpreted to be of palaeoproterozoic age. similar rocks which crop out south of sydostbugten to the north-east have previously been interpreted as along-strike equivalents of the naternaq supracrustal belt, although the intervening area is partly concealed by quaternary deposits (fig. 1). however, the discordant archaean granitic vein (483631) that cuts the metasedimentary rocks south of sydostbugten requires that the two belts contain rocks of different age. it is interesting to note that all the archaean detrital ages obtained match the ages between 2850 and 2800 ma of the archaean basement gneisses in the central nagssugtoqidian orogen. this distinguishes the metasedimentary rocks of the study area from metasedimentary rocks in the rinkian belt to the north, which include detrital zircons that are as old as 3600 ma (thrane et al. 2003). the predominance of archaean metasedimentary rocks unfortunately precludes them from being useful marker horizons to partition archaean and palaeoproterozoic deformation. nevertheless, the 2748 ± 19 ma synkinematic granite (463129) from saqqarput south-east of kangaatsiaq dates large fold structures in the northern nagssugtoqidian basement at around this age, which overlaps with the previously defined age of archaean deformation and metamorphism in the central part of the orogen (e.g. connelly & mengel 2000). the new age data place several constraints on the timing and style of palaeoproterozoic metamorphism and deformation in the region. the palaeoproterozoic sediment at naternaq (sample 464435) was probably metamorphosed and deformed shortly after its deposition, in line with the significant c. 1850 ma metamorphic peak in the archaean sediment from kangersuneq (sample 448394). the metamorphic zircon age from this sample 97 is not very precise, and tims analyses would be necessary to obtain an exact age of the metamorphism. however, it correlates with previous estimates for the timing of deformation and metamorphism both north and south of the study area (kalsbeek et al. 1987; kalsbeek & nutman 1996; whitehouse et al. 1998; connelly et al. 2000; thrane et al. 2003; connelly et al. 2006). furthermore, a late phase of n–s-directed shortening is dated at 1837 ± 12 ma by a synkinematic pegmatite (470515). the age of this pegmatite may correlate with the 1821 ma d2 deformation event defined in the central nagssugtoqidian orogen by connelly et al. (2000), and with 1821–1823 ma deformation east of ilulissat in the north (connelly et al. 2006). acknowledgements funding by the carlsberg foundation (thrane) and the national science foundation (grant ear-0337594 to connelly) is gratefully acknowledged. we thank ian alsop (university of st. andrews), adam a. garde and jeroen van gool (geus) for providing samples. we also thank john lansdown (the university of texas at austin) for assisting in collecting the la-icp-ms data, and julie a. hollis and adam a. garde (geus) for re-analysing sample 464435 zircons on the ion microprobe. allen p. nutman and åke johansson are thanked for their constructive comments, and adam a. garde and a.k. higgins are thanked for help with the geological introduction. references árting, u.e. 2004: a petrological study of basic dykes and sills of assumed palaeoproterozoic age in central western greenland, 122 pp., two appendices. unpublished m.sc. thesis, university of copenhagen, denmark. connelly, j.n. & mengel, f.c. 2000: evolution of archean components in the nagssugtoqidian orogen, west greenland. geological society of america bulletin 112, 747–763. connelly, j.n., van gool, j.a.m. & mengel, f.c. 2000: temporal evolution of a deeply eroded orogen: the nagssugtoqidian orogen, west greenland. canadian journal of earth sciences 37, 1121–1142. connelly, j.n., thrane, k., krawiec, a.w. & garde, a.a. 2006: linking the palaeoproterozoic nagssugtoqidian and rinkian orogens through the disko bugt region of west greenland. journal of the geological society (london) 163, 319–335. garde, a.a. 2004: geological map of greenland, 1:100 000, kangaatsiaq 68 v.1 syd. copenhagen: geological survey of denmark and greenland. glassley, w.e. & sørensen, k. 1980: constant ps-t amphibolite to granulite facies transition in agto (west greenland) metadolerites: implications and applications. journal of petrology 21, 69–105. hollis, j.a., keiding, m., stensgaard, b.m., van gool, j.a.m. & garde, a.a. 2006: evolution of neoarchaean supracrustal belts at the northern margin of the north atlantic craton, west greenland. in: garde, a.a. & kalsbeek, f. (eds): precambrian crustal evolution and cretaceous–palaeogene faulting in west greenland. geological survey of denmark and greenland bulletin 11, 9–31 (this volume). kalsbeek, f. & nutman, a.p. 1996: anatomy of the early proterozoic nagssugtoqidian orogen, west greenland, explored by reconnaissance shrimp u-pb dating. geology 24, 515–518. kalsbeek, f., pidgeon, r.t. & taylor, p.n. 1987: nagssugtoqidian mobile belt of west greenland: a cryptic 1850 ma suture between two archaean continents – chemical and isotopic evidence. earth and planetary science letters 85, 365–385. ludwig, k.r. 1999: isoplot/ex version 2.00 – a geochronological toolkit for microsoft excel. berkeley geochronology center, special publication 2. marker, m., whitehouse, m., scott, d., stecher, o., bridgwater, d. & van gool, j.a.m. 1999: deposition, provenance and tectonic setting for metasediments in the palaeoproterozoic nagssugtoqidian orogen, west greenland: a key for understanding crustal collision. abstracts eug 10, strasbourg, france. mazur, s., piazolo, s. & alsop, g.i. 2006: structural analysis of the northern nagssugtoqidian orogen, west greenland: an example of complex tectonic patterns in reworked high-grade metamorphic terrains. in: garde, a.a. & kalsbeek, f. (eds): precambrian crustal evolution and cretaceous–palaeogene faulting in west greenland. geological survey of denmark and greenland bulletin 11, 163– 178 (this volume). nutman, a.p., kalsbeek, f., marker, m., van gool, j.a.m. & bridgwater, d. 1999: u-pb zircon ages of kangâmiut dykes and detrital zircons in metasediments in the palaeoproterozoic nagssugtoqidian orogen (west greenland): clues to the pre-collisional history of the orogen. precambrian research 93, 87–104. østergaard, c., garde, a.a., nygaard, j., blomsterberg, j., nielsen, b.m., stendal, h. & thomas, c.w. 2002: the precambrian supracrustal rocks in the naternaq (lersletten) and ikamiut areas, central west greenland. geology of greenland survey bulletin 191, 24– 32. piazolo, s., alsop, g.i., van gool, j. & nielsen, b.m. 2004: using gis to unravel high strain patterns in high grade terranes: a case study of indentor tectonics from west greenland. in: alsop, g.i. et al. (eds): flow processes in faults and shear zones. geological society special publication (london) 224, 63–78. schuhmacher, m., de chambost, e., mckeegan, k.d., harrison, t.m. & migeon, h. 1994: in situ dating of zircon with the cameca ims 1270. in: benninghoven, a. (ed.): secondary ion mass spectrometry sims ix, 919–922. chichester: wiley. thrane, k., connelly, j.n., garde, a.a., grocott, j. & krawiec, a. 2003: linking the palaeoproterozoic rinkian and nagssugtoqidian belts of central w. greenland: implications of new u-pb and pb-pb ages. european union of geosciences meeting, geophysical research abstracts, cd 5, abstract # 09275. 98 van gool, j.a.m. et al. 2002: precambrian geology of the northern nagssugtoqidian orogen: mapping in the kangaatsiaq area. geology of greenland survey bulletin 191, 13–23. whitehouse, m.j., claesson, s., sunde, t. & vestin, j. 1997: ion microprobe u-pb zircon geochronology and correlation of archaean gneisses from the lewisian complex of gruinard, north-western scotland. geochimica et cosmochimica acta 61, 4429–4438. whitehouse, m.j., kalsbeek, f. & nutman, a.p. 1998: crustal growth and crustal recycling in the nagssugtoqidian orogen of west green__________________________________________________________________________________________________________________________________________________________________________________ manuscript received 15 october 2004; revision accepted 9 september 2005 land: constraints from radiogenic isotope systematics and u-pb zircon geochronology. precambrian research 91, 365–381. wiedenbeck, m., allé, p., corfu, f., griffin, w.l., meier, m., oberli, f., von quardt, a., roddick, j.c. & spiegel, w. 1995: three natural zircon standards for u-th-pb, lu-hf, trace element and ree analyses. geostandards newsletter 19, 1–23. williams, i.s. 1998: u-th-pb geochronology by ion microprobe. in: mckibben, m.a., shanks iii, w.c. & ridley, w.i. (eds): applications of microanalytical techniques to understanding mineralising processes. reviews in economic geology 7, 1–35. 99 appendix analytical methods rock samples were crushed to mineral size under clean conditions using a jaw crusher and a disc pulveriser, and initial mineral separation was made using a wilfley table at the university of copenhagen or the university of texas at austin. all subsequent procedures, including sieving, heavy liquids and magnetic separation, were conducted at the university of texas at austin. mineral fractions were characterised using a binocular reflected light microscope, a transmitted light petrographic microscope (with condenser lens inserted to minimise edge refraction), and a scanning cathodoluminescence (cl) imaging system on a jeol 730 scanning electron microscope. selected zircons of comparable size were hand picked and placed on two-sided tape, collared, and covered with epoxy. the zircons in the resulting mount were ground to approximately two thirds of their original thickness and polished. cl imaging was used to characterise the zircons before analysis. laser ablation analysis utilised a merchantek 213nm yag-laser connected to a micromass quadrupole mass spectrometer (platform). fractionation and inherent detector non-linearity were accounted for by analysing zircons already well characterised by tims. corrections necessary to obtain the correct 207pb/206pb ratios for these internal laboratory standards, covering a range of intensities and isotopic ratios, were applied to unknowns. standards were run throughout each session. blanks and offpeak baselines were also determined throughout each analytical session. selecting and analysing only the highest quality zircons minimised the need for common pb corrections using measured 204pb, a procedure made impossible by high 204hg counts. a single zircon analysis comprises approximately 450 ten-microsecond scans of 207pb/ 206pb. ratios reflecting the moving average of twenty 207pb and 206pb measurements are first plotted on a graph to check for anomalous ratios throughout a run. those with high ratios at the beginning are presumed to reflect common pb and are removed from further consideration. a jump from one ratio plateau to another during one analysis is interpreted to reflect piercing a core or rim of different age. only data from one plateau at a time were considered. in cases where the beam pierces the grain too deeply and ejecta are not effectively emitted towards the end of an analysis, the signal intensity and commonly also isotope ratios change dramatically. data from the late part of such runs were also rejected. with this first assessment of data complete, the 207pb and 206pb data were then passed through a 4-sigma filter to remove highly anomalous counts before being passed through a more rigorous 2-sigma filter. the averages of the remaining individual measurements (typically < 5% rejection) of 207pb and 206pb provided the final 207pb/206pb ratio and consequent age. given the transient signal inherent in la-icp-ms and sequential acquisition required by the single collector, standard statistics on multiple blocks of scans is not applicable. a major limitation of our la-icp-ms protocol is that u abundances are not measured; instead only 207pb/206pb ratios are obtained. a direct indication of concordance is, therefore, not available and all the ages obtained should conservatively be interpreted to represent minimum ages for crystallisation. a single sample was analysed using the cameca ims 1270 ion microprobe at the nordsim laboratory, naturhistoriska riksmuseet, stockholm. the sample was prepared in the same way as the samples analysed by laicp-ms. reference zircon 91500 from ontario, canada, with a weighted average 207pb/206pb age of 1065 ma (wiedenbeck et al. 1995), was included in the mount and used as standard. analytical procedures and common lead corrections are similar to those described by schuhmacher et al. (1994) and whitehouse et al. (1997). calibrations of pb/u ratios are based on the observed relationship between pb/u and uo2/u and follow procedures similar to those used by the shrimp group at the australian national university (williams 1998). 100 research article wehrlé et al. 2021: geus bulletin 47. 5284. https://doi.org/10.34194/geusb.v47.5284 1 of 9 greenland bare-ice albedo from promice automatic weather station measurements and sentinel-3 satellite observations adrien wehrlé*1 , jason e. box1 , masashi niwano2 , alexandre m. anesio3 , robert s. fausto1 1geological survey of denmark and greenland (geus), copenhagen, denmark, 2meteorological research institute, japan meteorological agency, tsukuba, japan, 3department of environmental science, aarhus university, roskilde, denmark abstract the programme for monitoring of the greenland ice sheet (promice) provides surface meteorological and glaciological measurements from widespread on-ice automatic weather stations since mid-2007. in this study, we use 105 promice ice-ablation time series to identify the timing of seasonal bare-ice onset preceded by snow cover conditions. from this  collection, we find a bare-ice albedo at ice-ablation onset (here called bare-iceonset albedo) of 0.565 ± 0.109 that has no apparent spatial dependence among 20 sites across greenland. we then apply this snow-to-ice albedo transition value to measure the variations in daily greenland bare-ice area in sentinel-3 optical satellite imagery covering the extremely low and high respective melt years of 2018 and 2019. daily greenland bare-ice area peaked at 153 489 km² in 2019, 1.9 times larger than in 2018 (80 220 km²), equating to 9.0% (in 2019) and 4.7% (in 2018) of the ice sheet area. 1 introduction the recent net loss of greenland land ice is among the largest contributors to global sea-level rise (box & sharp 2017). while warm air advection produces the highest daily ice-ablation observations (fausto et al. 2016), absorbed sunlight is the largest melt energy source over seasonal time scales (van den broeke et al. 2008; box et al. 2012; fausto et al. 2016). absorbed sunlight increases during the melt season as surface conditions shift from a highly reflective, dry snow cover, to lower albedo wet snow with larger grains ( wiscombe & warren 1980; brun 1989), and yet lower albedo across the ablation area. bare ice darkened primarily by ice algae (stibal et al. 2017; ryan et al. 2018; cook et al. 2020; williamson et al. 2020) plays an important role in peak ice-sheet melt rates. accurate definition of bare-ice albedo at ice-ablation onset (hereafter bare-ice-onset albedo) has applications in (1) classifying the bare-ice area over large areas of the ice sheet (ryan et al. 2019; fausto et al. 2020); (2) constraining polar regional climate models used to estimate the surface mass balance of the greenland ice sheet ( fettweis et al. 2020) and (3) climate monitoring (e.g. moon et al. 2020). here, we study the surface climate conditions spanning the melt season transition from dry snow to bare ice using promice ground measurements. our main objective is to determine an albedo value useful in classifying the boundary between seasonal snow cover and bare ice. we proceed to determine the spatial and temporal patterns of bare-ice albedo using spaceborne *correspondence: adrien.wehrle@hotmail.fr received: 09 jun 2020 accepted: 12 jan 2021 published: 19 apr 2021 keywords: greenland ice sheet, albedo, ice ablation, promice, sentinel-3 abbreviations: aws: automatic weather station promice: programme for monitoring of the greenland ice sheet olci: ocean and land color instrument cpi: cloud cover probability index toa: top of atmosphere sza: solar zenith angle scda: simple cloud detection algorithm slstr: sea and land surface temperature radiometer rmse: root-mean-square error modis: moderate resolution imaging spectroradiometer geus bulletin is an open access, peerreviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: signe hillerup larsen (geus, denmark) reviewed by: horst machguth (université de fribourg, switzerland), xavier fettweis (university of liege, belgium) funding: see page 8 competing interests: none declared additional files: see page 8 https://doi.org/10.34194/geusb.v47.5284 https://orcid.org/0000-0002-4870-1821 https://orcid.org/0000-0003-0052-8705 https://orcid.org/0000-0003-3121-3802 https://orcid.org/0000-0003-2990-4014 https://orcid.org/0000-0003-1317-8185 mailto:adrien.wehrle@hotmail.fr wehrlé et al. 2021: geus bulletin 47. 5284. https://doi.org/10.34194/geusb.v47.5284 2 of 9 www.geusbul let in.org observations from the eu copernicus sentinel-3 satellite mission applied to a low (2018) and high (2019) melt year to measure the maximum relative differences in bare-ice area. 2 data and methods 2.1 promice surface measurements across the greenland ice sheet, the sunlight reflectivity of snow and ice, hereafter albedo, and several other surface meteorological and glaciological parameters are measured by more than 20 automatic weather stations (awss), operated by the programme for monitoring of the greenland ice sheet (promice) since mid-2007 (ahlstrøm et al. 2008). promice operates awss in nine regions around greenland (fig. 1) where most locations have a lower and upper aws, denoted by _l and _u, respectively. in our analysis of daily average promice aws data, we used an initial dataset of 225 station years from 26 station locations for air temperature, snow thickness, ice ablation and albedo. three stations situated in the accumulation area, where the underlying ice did not appear, were excluded from the analysis. the acquisition and/or computation of each variable is described in the following sections. 2.1.1 air temperature air temperature is recorded at promice awss using a platinum resistance thermometer in an aspirated shield. the measurement height above the surface varies between 0 and 2.6 m due to snow accumulation, compaction and ablation. 2.1.2 seasonal snow layer thickness snow thickness above the ice surface is obtained from the aws acoustic recordings of distance from a sonic sensor to the snow or ice surface. the sensor height above bare ice is determined here for each station year from a 20-day average of daily values. this average is computed 10 days after bare-ice onset determined from ice-ablation measurements (see section 2.2) to ensure fig. 1 locations and description of programme for monitoring of the greenland ice sheet (promice) automatic weather stations (awss). a: promice sites. dashed lines indicate surface elevation in metres. red circles indicate promice awss used in this study while blue circles indicate awss excluded from this study. b: promice aws instruments. red circles indicate the instrumentation used in this study. 80°n 75°n 70°n 65°n 60°n 80°w 60°w 50°w 40°w 20°w 20°w 0°w 225 0 3000 2750 2500 100 0 150 0 175 0 20 0022 50 27 5025 00 22 50 17 50 1500 12 50 200 0 thu greenland 500 km greenland tas mit egp qas nuk upe kpc sco kan 1: solar and infrared radiation, 2: tilt sensor, 3: satellite antenna,4:wind speed & direction, 5:snow/ice surface height, 6: air temperature & humidity, 7: ice ablation ’hose’, 8: solar panel, 9: data logger, barometer and gps, 10: battery, 11: ice temperature profile (8 levels) 4 3 2 1 5a 6 7 8 5b 9 7 10 11 cen https://doi.org/10.34194/geusb.v47.5284 http://www.geusbulletin.org wehrlé et al. 2021: geus bulletin 47. 5284. https://doi.org/10.34194/geusb.v47.5284 3 of 9 www.geusbul let in.org snow-free conditions. the sensor height is then subtracted to obtain the snow thickness. 2.1.3 ice ablation promice stations measure ice ablation using the pressure of the column of antifreeze over a pressure transducer (fausto et al. 2012). the transducer is initially drilled 10 to 14 m into the ice and re-installed to avoid complete exhumation. at high ablation (>6 m ice per year) sites like qas_l and kan_l, instrument re-installation occurs each year. at most other sites, the re-installation frequency is two to three years. 2.1.4 surface albedo daily average albedo is calculated from 10-minute tilt corrected (wang et al. 2015) upward and downward solar irradiance recordings in the 0.3 to 2.5 μm wavelength range. hourly data are averaged for cases with solar elevation angles above 20°. daily averages are computed from hourly data between 0 and 1. the daily albedo values are further adjusted after the correction proposed by aoki et al. (2011) for measurement platform obstruction of the radiometer field of view. this correction increases promice aws albedo measurements by 0.034 on average. 2.2 determining greenland bare-ice albedo at ice-ablation onset in order to study the ablation season albedo as it transitions from snow to bare ice, we use ice-ablation measurements to determine the timing of bare-ice onset. seasonal snow layer thickness and air temperature further contributed to a better understanding of the evolution of surface conditions. ice-ablation time series were manually compensated for signal shifts caused by station movement, sensor reinstallation and measurement failure. we then attempted to automatically determine a theoretical candidate date of ice-ablation onset for each station year. to this end, an automatic detection was conducted to identify curve inflexions (satopaa et al. 2011). however, signal variability as well as heterogeneous transition patterns precluded this approach. instead, the ice-ablation onset was first identified manually for each station year with supporting data of snow layer thickness and air temperature time series when necessary. despite careful inspection, roughly 20% of the ice ablation onset dates were flagged as ambiguous because of noisy or complex patterns. also, distinguishing transient melting days before the ice melt onset from a short sequence of freezing days after the start of ice ablation remained challenging. measurement uncertainty was not taken into account in this initial and theoretical estimation of ice-ablation onset but helped to better constrain its identification before further refinements. following this first step, we excluded time series with interruptions within the period of interest, instrument malfunctions and/or those for which the ice ablation onset could not be identified. only data within ± 45 days of bare-ice onset were thereafter considered, corresponding to the average time span of the ablation season. the resulting dataset contained 105 station years from 20 stations across the ice sheet (c. 47 and 77% of the initial dataset size, respectively) consisting of more than 9000 daily measurements for each variable. the raw ice-ablation measurements, consisting of cumulated values after each instrument re-installation, the average instrument recording within 45 days before the initial bare-ice onset were then subtracted from the associated ice-ablation time series for each station year. in this way, the ablation measurements forming the final dataset consisted of values relative to the premelt season and therefore, each station year had a zero average prior to the ice-ablation season. the processing steps are further documented in a github repository (wehrlé & box 2020a). the timing of bare-ice onset was then refined for each station year by accounting for measurement uncertainty. to this end, a threshold was determined to estimate the first ‘significant’ ice-ablation value after the manually selected ablation onset; a value for which we have high confidence that bare-ice conditions begin to prevail. a too conservative ice ablation threshold (i.e. above the actual precision of the measurement), would lead to a delayed detection of the bare-ice onset date. such a delay would cause a dark bias in the determination of the average bareice albedo as the ice may have already been affected by algal darkening (stibal et al. 2017; ryan et al. 2018; cook et al. 2020; williamson et al. 2020). on the other hand, bare-ice onset would be defined prematurely if the threshold was too restrictive (i.e. below the actual precision of the measurement) leading to a bright bias in bare-ice albedo associated with residual patches of snow. assessment of the temporal variation in surface conditions spanning the melt season snow-to-ice transition was therefore needed to determine an optimal threshold. for this reason, we computed the average bare-ice albedo from all station years at the date of ice-ablation onset for a range of ice-ablation thresholds. the lower threshold boundary, set to 4 cm, corresponds to the accuracy of the pressure signal in the ice-ablation setup for each measurement. on the other hand, 8 cm which is the sum of the uncertainties of two pressure measurements, represents a maximized theoretical uncertainty of the ice-ablation https://doi.org/10.34194/geusb.v47.5284 http://www.geusbulletin.org wehrlé et al. 2021: geus bulletin 47. 5284. https://doi.org/10.34194/geusb.v47.5284 4 of 9 www.geusbul let in.org measurement. we added 10% of this value to obtain a conservative upper limit threshold of 9 cm. we found a curve inflection in the resulting albedo at a threshold of 6 cm with the slope of the linear regression being 11 times steeper within the 4–6 cm range than within the 6–9 cm range (–0.98 and –0.09, respectively). we contend that this change in data behaviour illustrates a snow-to-ice transition. as the melt of remaining heterogeneous patches of snow rapidly decreases the albedo, the underlying ice appears and further darkens at a slower rate. to be certain of bare-ice prevalence, we therefore defined the day of bare-ice onset as the first day after the manually selected transition with a cumulative ice ablation >6 cm. the 6 cm threshold equals the ice-ablation uncertainty determined in fausto et al. (2016). the resulting start of significant ice ablation occurs 4 ± 3 days after the manually selected inflection determined in the first step. the time series of each climate variable was then combined to build composites, that is, multi station-year averages synchronised to the determined emergence of bare ice. the average albedo value at the date of bare-ice onset, our bare-ice-onset albedo, was finally extracted. 3 results in this section, we present and analyse surface conditions ± 45 days around the onset of ice ablation with an emphasis on the albedo transition from snow to ice. 3.1 air temperature eight days prior to average bare-ice onset, composite average air temperature increases on average by 0.18°c per day before reaching the melting point (fig. 2a). after four days of relatively constant air temperatures under snow-melt conditions (0.21 ± 0.09°c), composite air temperature increases with an average rate of 0.57°c per day until bare-ice onset. composite air temperature then stabilises at 2.75 ± 0.43°c, which we suggest is associated with the sensible heat sink effect of the isothermal melting surface. 3.2 seasonal snow layer thickness composite snow height decreases at a rate of 7 ± 8 mm per day over the first 31 days, then gradually increases to 16 ± 7 mm per day until bare-ice onset (fig. 2b). the high variability before bare-ice onset is associated with differences in weather conditions between stations. 3.3 ice ablation before melt onset, average near-zero values of ice ablation have a standard deviation of 3.4 cm, associated with signal noise. average rate of ice ablation is 3.5 ± 0.5 cm per day after bare-ice onset (fig. 2c). 3.4 surface albedo the daily average albedo composite is stable (0.794 ± 0.008) until c. 15 days prior to bare-ice onset. then, average albedo declines by –0.008 ± 0.007 per day until 6 days prior to bare-ice onset (fig. 2d). the albedo decline rate then increases to –0.029 ± 0.009 per day until bare-ice onset. this steeper decline is partly driven by emergence of darker bare-ice patches and snow metamorphism. at the onset of snow melt, wet snow metamorphism (brun 1989) causes rapid grain growth, resulting in the reduction of near-infrared snow albedo (wiscombe & warren 1980; brun 1989). the composite bare-ice-onset albedo is 0.565 ± 0.109, which is between the recommended values for superimposed ice and clean ice in cuffey & paterson (2010). values remain stable for 8 days before albedo further declines at an average rate of –0.009 ± 0.004 per day for 10 days. finally, a slight decrease (–0.002 ± 0.005 per day) brings the composite albedo to its minimum daily average over the entire period (0.454 ± 0.140), 36 days after bare-ice onset. the mean difference of 0.111 between bare-ice onset and minimum albedo may be the result of ice algal growth (see stibal et al. 2017). the composite average albedo subsequently increases due to a temperature decrease as the melt season comes to its end and seasonal snowfall begins to accumulate. we further examined a cloud cover probability index (cpi) based on longwave downward irradiance and air temperature (van as 2011) to assess the influence of cloud cover on albedo. we find a daily standard deviation of 0.013 on composite albedo, albedo depending on the cpi threshold (from 0.1 to 0.9, in 0.01 steps). on the day of bare-ice onset, a standard deviation of 0.007 suggests little influence of cloud cover. the significant variability associated with the composite bare-ice-onset albedo (0.565 ± 0.109) is a result of a combination of data acquisition factors, method accuracies and regional variations. while the measurement accuracies are estimated, we cannot assess the method accuracy or the regional variations in bareice-onset albedo because of a lack of in situ measurements at field-verified dates of bare-ice onset and with a widespread spatial coverage. thus, the relative share of responsibility of each of these sources in the total variability remains unknown. nevertheless, in order to investigate the stability of this bare-ice-onset albedo value, we conducted 10  000 simulations where half the station years (53) were excluded randomly. we found a standard deviation of 0.011 on average bare-ice-onset albedo, which demonstrates https://doi.org/10.34194/geusb.v47.5284 http://www.geusbulletin.org wehrlé et al. 2021: geus bulletin 47. 5284. https://doi.org/10.34194/geusb.v47.5284 5 of 9 www.geusbul let in.org a low sensitivity to sample size. we also found a low variability within the range of ice-ablation thresholds from 4 to 9 cm (daily albedo standard deviation of 0.005 for the period of interest), boundary thresholds being associated with bare-ice-onset albedo values of 0.585 and 0.560, respectively. while a relatively large difference in bare-ice-onset albedo (0.02) is obtained for ice-ablation thresholds between only 4 and 6 cm, a very small difference (0.002) is obtained for a larger difference of 3 cm between ice-ablation thresholds of 6 and 9 cm. these results further support the change in data behaviour observed at an ice-ablation threshold of 6 cm presented in section 2.2, where the evolution in bare-ice albedo switches from a quickly decreasing to a slowly decreasing regime. finally, low correlations between average bare-ice albedo and station elevation, latitude and longitude (–0.07, 0.25 and 0.04, respectively) suggest no spatial dependence of the bare-ice-onset albedo. 4 application to spaceborne observations the bare-ice-onset albedo determined in the previous section can be used as an upper bound for ice albedo in order to monitor the evolution of the greenland bareice area throughout the melt season using spaceborne observations. here, we focus on 2018 and 2019 which are low and high melt years, respectively. 4.1 albedo retrieval from sentinel-3 observations the eu copernicus sentinel-3 satellite ocean and land color instrument (olci) provides 21 spectral bands from 400 nm in visible wavelengths to 1020 nm in the near-infrared, from october 2016 to present. here, we computed snow albedo (asnow) from olci observations using a fast atmospheric correction technique (kokhanovsky et al. 2018, 2020). because the extremely heterogeneous bare-ice surface conditions violate assumptions in kokhanovsky’s theory, we determined bare-ice albedo from a simple empirical approach. this approach consists of a fit between 4729 hourly promice albedo measurements and the nearest in time and space olci top of atmosphere (toa) reflectances spanning three years (2017–2019). in order to compute this fit, olci toa radiances were first converted to reflectances (r) normalised with the solar zenith angle (sza) by: r = olci toa radiance π / (cos(sza) s0) (1) where s0 is the toa solar irradiance measured aboard sentinel-3. –45 –35 –25 –15 –5 5 15 25 35 45 –45 –35 –25 –15 –5 5 15 25 35 45 –45 –35 –25 –15 –5 5 15 25 35 45–45 –35 –25 –15 –5 5 15 25 35 45 –12.5 –10.0 –7.5 –5.0 –2.5 0.0 2.5 5.0 ai rt em pe ra tu re ,° c –2.0 –1.5 –1.0 –0.5 0.0 ic e ab la tio n, m –0.2 0.0 0.2 0.4 0.6 0.8 1.0 sn ow he ig ht ,m a b c ± one standard deviation d bare-ice day bare-ice day bare-ice day bare-ice day 0.3 0.4 0.5 0.6 0.7 0.8 0.9 al be do ,u ni tle ss 0.565 ± 0.109 fig. 2 multi-site and multi-year composite surface conditions synchronised to bare-ice onset from ice ablation (black vertical dashed lines). a: air temperature. b: snow height. c: ice ablation. d: albedo where the red horizontal dashed line indicates the bare-ice-onset albedo. grey shading corresponds to ± one standard deviation around daily averages. https://doi.org/10.34194/geusb.v47.5284 http://www.geusbulletin.org wehrlé et al. 2021: geus bulletin 47. 5284. https://doi.org/10.34194/geusb.v47.5284 6 of 9 www.geusbul let in.org we then defined the bare-ice albedo (aice) from a fit to the average of four olci toa reflectances: aice = α (r400 nm+ r560 nm+ r865 nm + r1020 nm) / 4+ β (2) where α corresponds to the slope of the linear regression between olci toa and promice albedo measurements, and β is its intercept. computing an orthogonal distance regression, we found respective α and β coefficients of 1.003 and 0.058 associated with a high correlation of r = 0.899 and a standard error of 0.006. we subsequently built daily mosaics for greenland albedo using asnow combined with aice for asnow values below the average bare-ice-onset albedo (0.565) determined in this study. we found an average absolute difference of 0.078 between aice and asnow, ±10% around the bare-ice-onset albedo. clouds were detected and thereafter masked in sentinel-3 imagery using the simple cloud detection algorithm (scda) version 2.0 (metsämäki et al. 2015; wehrlé & box 2021). this algorithm consists of up to six tests on sea and land surface temperature radiometer (slstr) toa reflectances (550 and 1600 nm) and brightness temperatures (3.7, 11 and 12 μm). we further applied a temporal filter based on outlier detection modified after box et al. (2017) to remove remaining cloud artifacts, which would otherwise introduce abrupt temporal variations in the albedo time series. this processing consists of a 10-day rolling average applied to each pixel time series, for cases where the central value is within ± 15% of the window median. to validate the albedo retrievals, we compared the snow and ice olci-derived albedo with promice ground measurements. we first used the cpi to exclude albedo ground measurements acquired under cloudy conditions that we associated with a cpi >0.3. despite the ‘collocation problem’ of the large difference in footprint size between the ground station (c. 2 m × 2 m footprint) and the 1 km × 1 km olci pixel (see ryan et al. 2017), we found a high correlation (r = 0.885, n = 549), a root-mean-square error (rmse) of 0.079 and an insignificant mean bias (0.005) between olci-derived albedo and ground measurements. further, we found a mean absolute difference of 0.069 within ± one standard deviation around the bare-ice-onset albedo (0.565 ± 0.109). we expect that some part of the rmse is attributable to transient errors in the ground observations, such as water droplets or ice on the radiometer domes, specular reflections from the station or equipment shadowing on the surface or radiometer. the assessed albedo rmse can also increase due to undetected clouds affecting the spaceborne observations. finally, ‘gapless’ daily 1 km olci albedo grids were generated by updating pixel values when an area is considered to be cloud free. the last valid pixel value covering a given area then remains until a new valid value is determined. the different processing and filtering steps are documented in a github repository (wehrlé & box 2020b). 4.2 deriving greenland bare-ice area the bare-ice-onset albedo (0.565) was used to compute bare-ice area from ‘gapless’ daily albedo averages in 2018 (low melt year) and 2019 (high melt year). the 0.565 threshold represents a compromise between the 105 station years of the dataset as it is associated with variability in ice types linked to background impurity concentration and resulting albedo. while the promice network provides unprecedented access to crucial observations of conditions at the ice-sheet surface, the apparent limitations of in situ monitoring at very high resolution still prevent a precise study of these spatial variations using ground measurements. the standard deviation of 0.109 on the bare-ice-onset albedo is therefore associated with the spatial and temporal variability of the parameters measured at the different awss. this standard deviation probably overestimates the true variability of bare-ice-onset albedo across the entire ice sheet. the inevitable localised characteristics of field measurements often result in small sampling sizes relative to the area of the region of interest, which leads to a higher influence of outliers compared to larger datasets. on the other hand, we estimate that sampling size has little influence on the average bare-ice albedo itself (see section 3). consequently, we decided not to use this standard deviation as a measure of uncertainty for our thresholding to determine bare-ice area. nevertheless, to assess the sensitivity of the bare-ice area determination, we applied bare-ice albedo values obtained with theoretical ice-ablation thresholds of 4 and 9 cm (0.585 and 0.560, respectively) introduced in section 2. 4.3 variations in greenland bare-ice area the sentinel-3 satellite-derived albedo maps upscale our analysis and compare the near record high melt year, 2019 (tedesco & fettweis 2020), to the low melt year, 2018, across greenland (fig. 3). daily greenland ice albedo (including peripheral ice caps) was on average 0.030 (–3.7%) lower in 2019 for the melt season defined as 1 may to 15 september, reaching a maximum difference of 0.053 (–6.7%) on 3 august (fig. 3a). this maximum difference occurred after a 2019 high-melt event, which reduced the 2019 daily albedo to a minimum of 0.738, lower by 0.044 (–5.6%) than the 2018 minimum of 0.782. as estimated from the moderate resolution imaging spectroradiometer (modis) after box et al. (2017), albedo averaged over greenland land-ice was 0.817 and 0.777 from june to august 2018 and 2019, https://doi.org/10.34194/geusb.v47.5284 http://www.geusbulletin.org wehrlé et al. 2021: geus bulletin 47. 5284. https://doi.org/10.34194/geusb.v47.5284 7 of 9 www.geusbul let in.org respectively. this is 0.013 and 0.006 higher than the values determined in this study for the same time span (0.804 and 0.771). on average from 1 may to 15 september, daily bareice area was 47  329 km² in 2019 (fig. 3b), 4.5 times larger than in 2018, and reached a maximum difference of 102 943 km² (3.1 times larger than the previous year) on 3 august (fig. 3c). this deviation in ratios of average and maximum values between 2019 and 2018 is a result of the early 2019 melt onset, while 2018 bare-ice area remained near zero for the same period. a maximum daily bare-ice area of 153  489 km² occurred in 2019, 73 269 km² or 1.9 times larger than the 2018 maximum (80 220 km²). in 2019, the maximum daily bare-ice area corresponded to 9.0% of the greenland ice sheet, only 4.7% in 2018. maximum daily bare-ice area occurred 14 days earlier in 2019 than 2018 (3 and 17 august, respectively), while minimum albedo occurred 9 days earlier in 2019 than 2018 (4 and 13 august, respectively). using a watershed algorithm (van der walt et al. 2014) to determine the outer boundaries of the ice sheet, we identified ice caps as ice bodies separated from the ice sheet. while ice caps represent only c. 3.6% of the ice-sheet surface area, their inclusion increases 2019 and 2018 maximum daily bare-ice areas by 17.2% and 15.5%, respectively, increasing calculated bare-ice area ratios to 10.4% and 5.3%. despite the extremely low ice-ablation threshold of 4 cm, corresponding to the uncertainty of a single pressure measurement, we found maximum daily bare-ice areas equivalent to ± 11.4% (2018) and ± 7.0% (2019) of the values determined with the selected threshold of 6 cm. by applying the high ice-ablation threshold of 9 cm, we found differences equivalent to ± 2.0% and ± 2.8% (for 2018 and 2019, respectively) further supporting a relatively low ice-ablation threshold sensitivity. 5 conclusions promice aws time series of air temperature, snow height, ice ablation and albedo provide insights into snow and meteorological processes at the ablation-driven transition from seasonal snow to bare-ice surface conditions. we identified 6 cm as the first significant measured value of ice ablation through a sensitivity analysis, which matches the ice ablation uncertainty determined by fausto et al. (2016). by applying the 6 cm threshold to 05 /01 05 /15 06 /01 06 /15 07 /01 07 /15 08 /01 08 /15 09 /01 09 /15 month/day 0 25 50 100 125 20 19 m in us 20 18 ba re -ic e ar ea ,1 03 km 2 0 50 100 150 ba re -ic e ar ea ,1 03 km 2 0.74 0.76 0.78 0.80 0.82 0.84 al be do ,u ni tle ss 2018 july july 2018 july 2019 0.4 0.5 0.6 0.7 0.8 albedo a b c 2019 75 fig. 3 monthly averaged greenland snow and ice albedo from sentinel-3 olci data in july 2018 and 2019. inset figures: time series of greenland ice-sheet albedo and bare-ice area over 2018 and 2019 melt seasons (1 may and 15 september). a: daily area-averaged albedo. b: daily bare-ice area. blue shading corresponds to the range of bare-ice area computed from bare-ice albedo threshold values of 0.585 and 0.560, according to ice-ablation thresholds of 4 and 9 cm, respectively. the asymmetry of the blue shading (lower parts confounded with the main curve) is a result of the change in bare-ice albedo as a function of the ice-ablation threshold discussed in section 2.2. c: daily difference in bare-ice area between 2019 and 2018. grey shading corresponds to the range of bare-ice area deviations computed from the two bare-ice albedo threshold values used in c. grey vertical shading highlights the month of july. https://doi.org/10.34194/geusb.v47.5284 http://www.geusbulletin.org wehrlé et al. 2021: geus bulletin 47. 5284. https://doi.org/10.34194/geusb.v47.5284 8 of 9 www.geusbul let in.org identify the date of bare-ice onset for each station year in a semi-automatic procedure, we determined a bareice-onset albedo of 0.565 ± 0.109 for the greenland ice sheet. this value is between the recommended values for superimposed ice and clean ice reported by cuffey & paterson (2010). after bare-ice onset, we found a further albedo decrease of 0.111, which may be the result of ice-algal growth. average ice ablation was 3.5 ± 0.5 cm per day while average air temperature remained roughly constant, suggesting that stable air temperatures were associated with the sink of heat energy during surface melting. we found no dependence of bare-ice albedo on elevation, latitude or longitude, suggesting that the bare-ice-onset albedo determined here is representative for locations in between promice awss. we further combined snow albedo after kokhanovsky et al. (2018, 2020) with bare-ice albedo estimated from a fit between sentinel-3 olci toa reflectances and promice albedo data. in a cross-validation using 4729  daily promice observations, we confirmed a high correlation coefficient (0.885), a rmse of 0.079 and an insignificant average bias (0.005) with promice ground measurements. applying promice-derived bare-ice-onset albedo to the sentinel-3 imagery we produced quantitative mapping of albedo and bare-ice area variations. the maximum daily bare-ice area was 1.9 times larger in 2019 than in 2018 (153  489 and 80  220 km², respectively), covering 9.0% and 4.7% of the ice sheet. peripheral ice caps increase bare-ice area estimates by 17.2% (2019) and 15.5% (2018). thus, the combination of ground and spaceborne observations yields powerful quantitative constraint on snow-cover dynamics across greenland ice. acknowledgements reviewers are thanked for constructive commentary. funding statement promice has been supported by the danish ministry of climate, energy and utilities. the work with sentinel-3 is part of the preoperational sentinel-3 snow and ice products (sice) project supported by european space agency (esa) contract 4000125043-esa/ao/19101/17/i-nb eo science for society. competing interests the authors declare no competing interests. author contributions aw: writing – original draft preparation, methodology, software. jb: conceptualization, methodology, visualization, writing – reviewing and editing. rf, mn, aa: writing – reviewing and editing. additional files additional files are available at https://doi.org/10.5281/zenodo.4244909, https://doi.org/10.5281/zenodo.4244905. data produced in this study are available at https://doi.org/10.22008/fk2/snx8l6. references ahlstrøm, a.p. et al. 2008: a new programme for monitoring the mass loss of the greenland ice sheet. geological survey of denmark and greenland bulletin 15, 61–64. https://doi.org/10.34194/geusb. v15.5045 aoki, t., kuchiki, k., niwano, m., kodama, y., 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https://doi.org/10.1109/icdcsw.2011.20 https://doi.org/10.1002/2017gl075958 https://doi.org/10.5194/tc-14-1209-2020 https://doi.org/10.5194/tc-14-1209-2020 https://doi.og/10.3189/002214311796405898 https://doi.og/10.3189/002214311796405898 https://doi.org/10.5194/tc-2-179-2008 https://doi.org/10.5194/tc-2-179-2008 https://doi.org/10.7717/peerj.453 https://doi.org/10.7717/peerj.453 https://doi.org/10.5281/zenodo.4244909 https://doi.org/10.5281/zenodo.4244905 https://doi.org/10.5281/zenodo.4244905 https://doi.org/10.22008/fk2/n0xwsj https://doi.org/10.22008/fk2/n0xwsj https://doi.org/10.1073/pnas.1918412117 https://doi.org/10.1073/pnas.1918412117 https://doi.org/2.0.co;2”>10.1175/1520-0469(1980)037<2712:amftsa>2.0.co;2 https://doi.org/2.0.co;2”>10.1175/1520-0469(1980)037<2712:amftsa>2.0.co;2 geological survey of denmark and greenland bulletin 15, 2008, 1-7 geological survey of denmark and greenland ministry of climate and energy geological survey of denmark and greenland bulletin 15 • 2008 review of survey activities 2007 edited by ole bennike and a.k. higgins geological survey of denmark and greenland bulletin 15 keywords geological survey of denmark and greenland, survey organisations, current research, denmark, greenland. cover photographs from left to right 1. field work in west greenland. photo: jakob lautrup. 2. field work in the mo-clay region of denmark. photo: stig a. schack pedersen. 3. wind mills in denmark. geus has been mapping the geology of new offshore wind mill farm areas. photo: jørn bo jensen. 4. launching of seismic equipment from the swedish icebreaker oden during the lomrog (lomonosov ridge off greenland) cruise to the arctic ocean in 2007. photo: martin jakobsson, stockholm university. frontispiece: facing page as part of the danish continental shelf project (http://a76.dk) bathymetric, seismic and gravimetric data were acquired in august and september 2007 during the lomrog cruise in the arctic ocean north of greenland. the lomrog project is a co-operation between institutions in denmark, sweden and other nations. results from the lomrog cruise will be published in forthcoming years. photo: martin jakobsson, stockholm university. chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, department of earth sciences, university of aarhus; minik rosing, geological museum, university of copenhagen; finn surlyk, department of geography and geology, university of copenhagen scientific editors: ole bennike and a.k. higgins editorial secretaries: jane holst and esben w. glendal referees: (dk = denmark; numbers refer to first page of reviewed article): troels aagaard, dk (85); anonymous (41, 45, 53, 61, 89); geo arnold, holland (81); niels balling, dk (53); björn e. berglund, sweden (57); morten bjerager, dk (13); lars ole boldreel, dk (37); ilke borowski, germany (81); jens bruun-petersen, dk (37); bjørn buchardt, dk (13, 89); ashton embry, canada (65); ida fabricius, dk (17); clark friend, uk (73); henrik friis, dk (9); thomas frisch, canada (77); ole graversen, dk (25); bent hasholt, dk (61); claus heilmann-clausen, dk (29); claus heinberg, dk (29); julie a. hollis, australia (69, 73); mads huuse, uk (33); rasmus jacobsen, dk (45); kurt kjær, dk (33); jens konnerup-madsen, dk (69); john a. korstgård, dk (25); lars kristiansen, dk (21); nicolaj krog larsen, sweden (41, 49); ole bjørslev nielsen, dk (21); henrik olsen, dk (49); ian parsons, uk (9); john s. peel, sweden (77); heikki seppä, finland (57); martin sønderholm, dk (65, 93); ole v. vejbæk, dk (17); jim williamson, uk (93). (page nos corrected aug. 2009) illustrations: jette halskov lay-out and graphic production: annabeth andersen printers: schultz grafisk, albertslund, denmark manuscripts submitted: 4 february – 17 march 2008 final versions approved: 28 may 2008 printed: 10 july 2008 issn 1603-9769 (review of survey activities) issn 1604-8156 (geological survey of denmark and greenland bulletin) isbn 978-87-7871-213-4 citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 15, 96 pp. available from geological survey of denmark and greenland (geus) • øster voldgade 10 • dk-1350 copenhagen k • denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk © de nationale geologiske undersøgelser for danmark og grønland (geus), 2008 for the full text of the geus copyright clause, please refer to www.geus.dk/publications/bull 4 7. review of survey activities 2007 f.g. christiansen 9. diagenesis influencing the porosity of upper jurassic reservoir sandstones, danish north sea r. weibel and n. keulen 13. correlation of carbon isotope events in the danish upper cretaceous chalk n.h. schovsbo, s.l. rasmussen, e. sheldon and l. stemmerik 17. geophysical imaging of porosity variations in the danish north sea chalk t. abramovitz 21. palaeogene diatomite deposits in denmark: geological investigations and applied aspects s.a.s. pedersen 25. cenozoic palaeogeography and isochores predating the neogene exhumation of the eastern north sea basin p. japsen, e.s. rasmussen, p.f. green, l.h. nielsen and t. bidstrup 29. a new neogene biostratigraphy for denmark k. dybkjær and s. piasecki 33. mapping of buried tunnel valleys in denmark: new perspectives for the interpretation of the quaternary succession f. jørgensen and p.b.e. sandersen 37. base quaternary in the danish parts of the north sea and skagerrak t. nielsen, a. mathiesen and m. bryde-auken 41. geology of outer horns rev, danish north sea j.b. jensen, p. gravesen and s. lomholt ghana uganda tanzania mozambique latvia greenland canada uk norway faroe islands kenya estonia spain poland chile germany sw ed en the netherlands denmark lithuania brazil madagascar malawi zambia yemen geus working areas 2007. orange areas are covered in this volume. for further information on other working areas please refer to our website: www.geus.dk/international 5 45. use of geochemistry in groundwater vulnerability mapping in denmark b. hansen and l. thorling 49. sedimentary facies and architecture of the holocene to recent rømø barrier island in the danish wadden sea p.n. johannessen, l.h. nielsen, l. nielsen, i. møller, m. pejrup, t.j. andersen, j. korshøj, b. larsen and s. piasecki 53. evidence of stretching of the lithosphere under denmark s. gregersen, l.v. nielsen and p. voss 57. environmental response to the cold climate event 8200 years ago as recorded at højby sø, denmark p. rasmussen, m.u. hede, n. noe-nygaard, a.l. clarke and r.d. vinebrooke 61. a new programme for monitoring the mass loss of the greenland ice sheet a.p. ahlstrøm and the promice project team 65. the north-east baffin bay region, offshore greenland – a new frontier petroleum exploration region u. gregersen 69. geochemistry of greenstones in the tasiusarsuaq terrane, southern west greenland a. scherstén, h. stendal and t. næraa 73. new zircon ages from the tasiusarsuaq terrane, southern west greenland t. næraa and a. scherstén 77. hans ø, celebrated island of nares strait between greenland and canada: from dog-sledge to satellite mapping p.r. dawes and t. tukiainen 81. from science to practice in implementing the european union’s water framework directive l.f. jørgensen, j.c. refsgaard and a.l. højberg 85. kensea – tsunami damage modelling for coastal areas of kenya j. tychsen, o. geertz-hansen and f. schjøth 89. laser ablation analysis of bivalve shells – archives of environmental information m.h. klünder, d. hippler, r. witbaard and d. frei 93. fully automated analysis of grain chemistry, size and morphology by ccsem: examples from cement production and diamond exploration n. keulen, d. frei, s. bernstein, m.t. hutchison, c. knudsen and l. jensen thailand laos pdr vietnam mongolia phillipines cambodia 7 after a very critical period with a major reorganisation of the scientific environment in denmark, 2007 created new stability for the geological survey of denmark and greenland (geus). a new law describing the role of geus was passed by the danish parliament in the spring, followed by changes in top management and a newly formalised collaboration between geus, the university of copenhagen and the uni ver sity of aarhus – the so-called geocenter denmark. this collaboration is very promising and will provide a much better chance of integrating young scientists in research activities. this fifth annual issue describes selected activities that geus and its partners carry out in denmark, greenland and internationally. together with the previous four issues it provides a good overview of the many different types of activities and projects, the advisory capacity and analytical services available from geus. this 2007 review contains a total of 22 four-page papers, 13 on denmark, five on greenland and four with the focus on methodology or international work. compared to the previous four that were biased towards greenland, this issue has a majority of papers related to activities in denmark: oil and gas, deep crustal structure, offshore wind farms, coastal processes in the danish wadden sea (vadehavet), groundwater vulnerability, raw materials, neogene and quaternary history, and palaeoclimate. chalk is the main key to the good national economy of denmark. better understanding of the depositional and diagenetic processes that formed the chalk and the petrophysical properties of this prolific rock is necessary to exploit the petroleum potential offshore. chalk is covered in two papers in this issue, while two other papers on diagenesis and ex humation are also important for understanding the petroleum potential of denmark. the deep crustal structure under denmark is presented in one paper and danish raw materials are described in a paper on the diatomite deposits of the palaeogene fur formation. today climate changes and resultant adaptation are in focus. several papers are important in this context, especially on coastal processes, understanding of the neogene and qua ternary history, variation in palaeoclimate and monitoring of the ice sheet in greenland. a dramatic increase in oil and metal prices in recent years – together with abundant new targets for exploration – is the background for the boom within the oil and mining industries in greenland. the number of licences – and the area covered by them – is historically large. previous and ongoing research and mapping activities together with an easy access to critical data at geus have been very important in the process of attracting industry to greenland. new possibilities within both the oil and mineral industries are described in papers on the sedimentary basins of baffin bay and on the basement terrains in south-western greenland. the research history of the disputed island hans ø in nares strait between greenland and canada is also outlined in one article. two papers describe the development of new analytical techniques, one on the automatic use of computer-controlled scanning electron microscopy (ccsem) very important in exploration and for quality control of industrial materials; and the other one on the use of laser ablation inductively coupled plasma mass spectrometry (la-icp-ms) for environmental studies of bivalves. geus is active in international collaboration through the eu and has ongoing activities in many different countries, especially in south-east asia and africa. in this issue there is one paper on tsunami damage modelling in kenya and another on implementing the eu’s water framework directive. in the summer of 2007 geus in collaboration with other danish and swedish institutions and the swedish polar research secretariat acquired data in the area north of green land as part of the danish continental shelf project (see frontispiece). this project aims to acquire the necessary technical data to support a danish claim for extending the juridical continental shelf beyond 200 nautical miles before 2014. review of survey activities 2007 flemming g. christiansen deputy director © geus, 2008. geological survey of denmark and greenland bulletin 15, 7 only. available at: www.geus.dk/publications/bull research article pedersen & gravesen 2021: geus bulletin 47. 8304. https://doi.org/10.34194/geusb.v47.8304 1 of 8 the geological framework for hvideklint, south-east denmark, using glaciodynamic sequence stratigraphy stig a. schack pedersen* , peter gravesen geological survey of denmark and greenland (geus), copenhagen, denmark abstract glaciodynamic sequence stratigraphy provides a practical model for grouping and classifying complex geological data to aid interpretation of past climatic and environmental development in quaternary successions. the principles of glaciodynamic sequence stratigraphy are applied here to summarise the complex glacial geological framework of hvideklint on the island of møn, south-east denmark. the framework of the superimposed deformed hvideklint is presented in a reconstructed geological cross-section of hvideklint. for the construction of the architecture of the glaciotectonic complex, the interpretation of structures below sea level was based on a detailed new survey of the cliff section combined with construction of  successive approximation balanced cross-sections. the new description is supported by drill hole data from the jupiter database. where chalk is not glaciotectonically deformed, the constructed depth to the top-chalk-surface is generally located about 30 m below sea level. in hvideklint, thrust sheets with chalk are exposed 20 m above sea level, and the balanced cross-section constructions indicate that the décollement surface for a hvideklint glaciotectonic complex is located about 80 m below sea level. between the décollement level and the top of the complex, two or more thrust-fault flat-levels and connecting ramps add to the complex architecture of hvideklint. 1 introduction the danish land surface is comprised of a complex sequence of deformed glacial deposits. the principles of glaciodynamic sequence stratigraphy provide a practical model for grouping and classifying such complex geological data to aid interpretation of past climatic and environmental development in quaternary successions. furthermore, the concept of glaciodynamic sequence stratigraphy presents a system for predicting the successions expected in drill holes or excavations during constructive works. here, we apply the principles of glaciodynamic sequence stratigraphy to summarise the complex glacial geological framework of hvideklint on the island of møn, south-east denmark (fig. 1) and present a new geological cross-section of hvideklint (fig. 2). this work is part of a larger project  at  the  geological survey of denmark and greenland (geus) to *correspondence: sasp@geus.dk received: 12 mar 2019 accepted: 2 nov 2021 published: 14 dec 2021 keywords: glaciotectonics, glacioldynamics, sequence stratigraphy, lower maastrichtian chalk abbreviations: geus: geological survey of denmark and greenland hk: hvideklint geus bulletin is an open access, peerreviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: catherine n jex (geus, denmark) reviewed by: daniel paul le heron (university of vienna, austria) and one anonymous reviewer funding: see page 7 competing interests: none declared additional files: none provided https://doi.org/10.34194/geusb.v47.8304 https://orcid.org/0000-0002-7867-5118 mailto:sasp@geus.dk pedersen & gravesen 2021: geus bulletin 47. 8304. https://doi.org/10.34194/geusb.v47.8304 2 of 8 www.geusbul let in.org produce  a  systematic geological description of the island (pedersen & gravesen 2021). the basic challenge in stratigraphy is to correlate successions from one observation point to another. some boundaries can be traced as lithological surfaces, from which we construct a layer-cake model of stratigraphy across a region. however, complex successions containing superimposed deformation structures or erosional irregularities may require a different approach. here, a dynamic model is required to construct a geological framework and explain the relationship between interpreted deposits and their boundaries. sequence stratigraphy is an established model for describing complex sedimentary systems. it explains the stacking patterns and origins of sedimentary successions driven by changing sea levels. during ice ages, sea-level change is predominantly controlled by the accumulation of ice in terrestrial ice sheets. m øns klint møn hvideklin t hjelm bugt 50002500 n 0 meters 50 km møn sweden germany denmark fig. 1 location of hvideklint, møn, south-east denmark. the position of the cross-section (fig. 2) is marked by a red arrow. contour intervals mark the depth of the pre-quaternary surface (the top of the cretaceous chalk). the construction of contours is based on an arcgis algorithm applied to the well data from møn (jupiter database, geus). the top-chalk surface occurs mainly at 20–30 m b.s.l. the exception of this is the areas where it is elevated above sea level in the glaciotectonic complexes. the contour map is simplified from pedersen & gravesen (2021). tectonic brecciation thrust fault lolland till formation mid danish till formation møns klint formation klintholm till formation kraneled formation 0 20 m -20 -40 -60 -80 hk4 hk3 hk2 hk1 ne 100 m hk11 hk13 hk12 hk10 hk9 hk8 hk7 hk6 hk5 sw 0 20 m -20 -40 -60 -80 hk5 hk4 stubberup have formation kobbelgård formation fig. 2 new geological cross-section of hvideklint (hk). the interpretation of structures below sea level is based on successive balanced cross-section constructions according to the concept described by pedersen (2005). chalk thrust sheets are annotated hk1 to hk13 from ne to sw. slightly modified from pedersen & gravesen (2021). https://doi.org/10.34194/geusb.v47.8304 http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 47. 8304. https://doi.org/10.34194/geusb.v47.8304 3 of 8 www.geusbul let in.org glaciodynamic sequence stratigraphy establishes a model for glacial sediments and their relationship to glacial tectonics (pedersen 2012). as an ice sheet builds up, a sequence of glaciofluvial sediments is deposited in the proglacial area, in front of the ice margin. the deposition of these sediments ended by the transgression of the ice sheet and deposition of a lodgement till along the base of the ice. the deposits below the till will be affected by glaciotectonic processes, primarily as a proglacial thrust and fold complex and secondly by creating an unconformity between the glaciotectonic complex and a subglacially deposited basal till. the glaciotectonic unconformity truncates the thrust and fold structures, above which a glacitectonite may develop by shearing along the base of the ice in the so-called deformational layer (pedersen 2012). 2 methods a detailed new survey of the cliff section was carried out in 2015–2016. combined with construction of successive approximation-balanced cross-sections, we constructed the architecture of the glaciotectonic complex and interpretated the structures below sea level. the new description of hvideklint is supported by previous descriptions of the cliff section (aber 1979 and an unpublished master thesis by tove krogh stockmarr 1996, see acknowledgement). it is further supported by drill hole data from the geus jupiter database and the constructed depth to the top-chalk-surface formerly presented in the cross-section e–f in the geological map 1:50 000 of møn (pedersen & gravesen 2006, 2021). we refer the reader to pedersen & gravesen (2021) for a description of methods. 3 hvideklint geological description hvideklint is a more than 1 km long, 20 m high coastal cliff, facing the western part of hjelm bugt on the island of møn, south-east denmark (fig. 1). the cross-section along the costal cliff strikes ne–sw (fig. 2), almost perpendicular to the main direction of the strike of thrust faults and direction of the fold axes (fig. 3). the cliff comprises more than 13 thrust sheets of upper cretaceous chalk overlain by quaternary glacial deposits. displacement of the glaciotectonic thrust sheets occurred in the later part of the weichselian glaciation (aber 1979). saalian, eemian and the oldest weichselian deposits are displaced along the cliff. the youngest weichselian deposits discordantly truncate the thrust fault complex. the geological framework of hvideklint consists of three main elements: (1) cretaceous chalk sits beneath a brecciated unconformity, overlain mainly by glaciolacustrine sediments, thrust up by the proglacial push of the swedish ice advance from the north-east; (2) the superimposition of the glaciotectonic complex by the northward push of the baltic ice advance and (3) heavily n a " n b n = n = fig. 3 stereogrammetric projections of structural orientations measured along the hvideklint section. stereograms are lower hemisphere, equal area (schmidt) net projections. contour intervals are semistatistic densities 1, 5, 10 and 15%. blue dotted lines indicate principal directions of stress/strain. a: fold axes orientations. black dots are fold axes representing the deformation from the north-east (the hvideklint glaciodynamic event, ice advance indicated by blue arrow). black triangles indicate second phase of deformation from the south (møns klint glaciodynamic event indicated by red arrow) and/or superimposed deformed fold axes. b: bedding planes (black dots) and thrust planes (black triangles) are plotted by their normal to the planes. main orientation of planes shows a concentration of strike and dip around 135°/20° ne. red triangles represent thrusting from the south. the two arrows represent similar glaciodynamic impacts as in a. https://doi.org/10.34194/geusb.v47.8304 http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 47. 8304. https://doi.org/10.34194/geusb.v47.8304 4 of 8 www.geusbul let in.org brecciated chalk-clay sediments and formation of mud diapirism (stockmarr 1996). 3.1 lithostratigraphy the lithologies present at hvideklint are summarised in figs 2 and 4 and described below. 3.1.1 cretaceous chalk the cretaceous chalk is a white, grey-white or slightly yellowish coccolith chalk. black chert (flint) nodules are common (fig. 5), but continuous bands of chert are absent. in some places, chalk and marl lithologies occur with conglomeratic features (fig. 6). these lithologies correlate to the uppermost campanian – lowermost maastrichtian (c. 72 ma), the transition from the boesdal member in the mandehoved formation to the hvidskud member in the lowermost part of the møns klint formation (thomsen 1995; surlyk et al. 2013). 3.1.2 quaternary deposits the oldest quaternary deposit at hvideklint is interglacial, marine clay of eemian age resting on saalian glacial deposits (fig. 6). mollusc fauna of eemian age occurs in the black to dark greyish clay in the lowest part of the quaternary lithostratigraphy. however, the eemian clay rests on an irregular unconformity above the chalk. the top of the chalk is often brecciated and forms a glacitectonite, and residual gravel preserved in pockets along the unconformity indicates the remnants of glacial deposits and glaciotectonics related to the glacial dynamics prior to eemian, here interpreted as saalian. the oldest till bed is the ristinge till formation (houmark-nielsen 1987), a red-brown diamictite deposited during ice advance across the baltic in the early– middle weichselian, often referred to as the old baltic ice. the reddish colour is generally interpreted as the chalk till glaciolacustrian clay ny borre fm mid danish till fm stubberup have fm kobbelgård fm klintholmtill fm kraneled fm ristinge till fm eemian saalian upper cretaceous w ei ch se lia n lolland till fm fig. 4 simplified geological succession at hvideklint. geological units as in fig. 2. fig. 5 upper cretaceous chalk at hvideklint. note the occurrence of nodular flint in the chalk. fig. 6 brecciated chalk interpreted here as a mixture of conglomeratic chalk and marl at the boundary between campanian and maastrichtian chalk. the shear structures indicate superimposed deformation by shearing along the thrust fault at the base of the chalk. the unit is classified as a chalk-tectonite. https://doi.org/10.34194/geusb.v47.8304 http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 47. 8304. https://doi.org/10.34194/geusb.v47.8304 5 of 8 www.geusbul let in.org debris plugged up during the ice advance over the pre cambrian nexø sandstone and similar lithologies in the eastern baltic. the ristinge till formation is overlain by a heterolithic unit of glaciofluvial and glaciolacustrine sediments named the kraneled formation (fig. 4), which is capped by the dark grey coloured klintholm till formation (houmark-nielsen 1994, 2010; pedersen & gravesen 2021). these units are mainly shear-mixed and difficult to differentiate in detail. the klintholm till formation was previously dated to the middle weichselian, c. 34–30 ka bp (houmark-nielsen 2010; houmark nielsen et al. 2016). above the klintholm till formation is a claydominated unit of glacial deposits, referred to as the kobbelgård formation (pedersen & gravesen 2021; fig. 4). this unit is thought to have formed in a large lake covering the baltic depression during the interstadial period prior to the late glacial maximum in the last part of the weichselian (pedersen & gravesen 2009). in the top of the kobbelgård formation, deposits of glaciofluvial sand are intercalated in the clay. at the ne part of møn, the glaciofluvial sand unit is defined as the stubberup have formation (fig. 6; pedersen & gravesen 2021). below the thrust zones, sandy lithologies are almost absent at møns klint, where the clay in the kobbelgård formation (formerly named elephant clay) typically occurs below the base of the chalk thrust sheets. the mid danish till formation, which truncates the tips of the thrust sheet, is a grey-brownish, basal till, representing the last glacial maximum during the final stadial in the weichselian (houmark-nielsen 1987). it was deposited during the advance of the swedish ice stream shortly after the hvideklint glaciotectonic complex formed in the proglacial area. the glacial succession at hvideklint is capped by an orange-yellow-brown till. below the base of this till, the ny borre formation is located in other parts of møn (pedersen & gravesen 2021), but at hvideklint, it is only represented by an unconformity between mid danish till and the top till. the latter is the lolland till formation, deposited by the young baltic ice advance about 18 ka bp (pedersen et al. 2015). the young baltic ice advanced northwards from baltic, where it terminated at its stationary line in the central part of kattegat (pedersen & petersen 1997). the interval between the mid danish till formation and the lolland till formation is separated by a thin, irregular unit of sand and gravel, indicating that the swedish ice stream had melted back before the young baltic ice transgressed the area. 3.2 structural geology the hvideklint cross-section comprises more than 13 thrust sheets of chalk, piggybacked by a unit of glacial deposits (hk1–13, fig. 2; hk: hvideklint). based on an arcgis construction of the well data from møn (jupiter database, geus), the surface of the top of the chalk (pre-quaternary unconformity) occurs mainly at 20–30 m b.s.l., except where it is elevated above sea level in the glaciotectonic complexes (fig. 1). after drafting the cross-section of the coastal cliff section at hvideklint, balanced cross-sections were constructed following the principles of line-balance and displacement off-set. the base level for the displaced chalk thrust-sheets, defined as the décollement surface, was subsequently located 70–80 m b.s.l. the construction of the structures below sea level is a central new contribution, which is not integrated in the cross-section provided by aber (1979). the thrust sheets are annotated hk1–hk13 from ne to the sw in fig. 2. the outcrop of the westernmost thrust sheet (hk13) and thrust sheets west of this are at present obscured by screes (west of the cross-section depicted in fig. 2), mirroring the dominance of clayey units in the coast cliff. stereogrammetric plots of structural orientations measured along the hvideklint section are shown in fig. 3. in general, the thrust faults strike se–nw (c. 135°) and the dips of thrust fault ramps vary from 45° ne at the base of the cliff, becoming westerly (bending towards sub-horizontal) at the top of the cliff section (fig. 3). the thrust fault complex dominated by the glaciotectonic deformation from the north-east is superimposed by deformation from the south. this is best illustrated by the double-fold on the back of thrust sheet hk3, just in front of the tip of hk2 (fig. 2). here, the kobbelgård formation is tightly folded with inversion of the southern limb in a syncline fold with a southerly dipping axial plane (fig. 7). this type of superimposed structural feature, combined with an oblique cut of the structures (cross-section cut is not perpendicular to the direction of fold axis), is responsible for most of the irregularities identified in the cliff section. shear deformation with stretching of chalk debris and thin chalk flows is another important disturbance in the complex (fig. 8). a tectonic breccia formed the base of the thrust sheet, where the hanging-wall thrust sheet is displaced, and part of the décollement zone is thrust up along the ramp. here, the tectonically brecciated chalk is regarded as a tectonite. another type of tectonite formed along the base of the glacier ice, where the deformed layer sheared the top of the thrust sheets. the chalk breccias formed here are classed as chalk-glacitectonites (pedersen 1988). in the ideal glaciodynamic succession, the glacitectonites should always occur beneath a basal till, indicating that the first phase of ice transgression involved strong shearing along the base of the ice, after which basal till accumulated. https://doi.org/10.34194/geusb.v47.8304 http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 47. 8304. https://doi.org/10.34194/geusb.v47.8304 6 of 8 www.geusbul let in.org 3.3 glaciodynamic development a glaciodynamic sequence corresponds to a glaciodynamic event (pedersen 2012). at the base of the sequence, a décollement zone develops during the formation of a tectonite. dynamically, this develops simultaneously with the deposition of proglacial meltwater sand and clay (fig. 9, phase a). the second phase in the glaciodynamic event displaces the thrust sheets comprising the pre-quaternary units overlain by the proglacial deposits (fig. 9, phase b). during the third phase, the thrust fault complex is transgressed by the ice, which truncates the top of the complex, establishing a glaciotectonic unconformity under the formation of a glacitectonite (fig. 9, phase c). the basal till above the glacitectonite marks the end of the glaciodynamic sequence (fig. 9, phase d). some sequences may only be partially represented, due to poorly preserved glacial deposits or underdeveloped glaciotectonic deformation (fig. 9). at hvideklint, we have established four glaciodynamic events: (1) the saalian glaciation, (2) the klintholm glaciodynamic event, (3) the hvideklint glaciodynamic event, and (4) the møns klint glaciodynamic event. an early glaciodynamic event of saalian (or older) age is represented by the glaciotectonic brecciations of the top of the chalk, upon which the eemian clay was a b fig. 7 refolded beds of hvideklint units. a: the position of the overturned syncline in the hvideklint coastal cliff. b: close-up of the overturned syncline with the heterolithic beds of the kobbelgård and stubberup have formations. fig. 8 chalk-tectonite formed along the décollement zone below the displaced thrust sheet of hk4 (see fig. 2). note that a tectonite is the tectonic breccia below the displaced thrust sheet, in contrast to a glacitectonite, which is the tectonic breccia formed underneath the inferred former ice base. d basal till glacitectonite thrust faults sand and clay décollement cretaceous chalk c b a fig. 9 the glaciodynamic sequence system as applied to hvideklint. phase a: the décollement zone consists of a tectonite at the base of the thrust sheets, which may include mobilised mud diapirs. glacio-fluvial and glacio-lacustrine units are deposited prior to proglacial deformation. phase b: an allochthonous unit of pre-quaternary and/or pre-glacial sedimentary deposits, bounded by thrust faults representing the second order surfaces in the glaciotectonic architecture. phase c: a glaciotectonic unconformity including a glacitectonite is formed at the base of the advancing ice. phase d: the glaciodynamic sequence is terminated by the deposition of the basal till. slightly modified from pedersen & gravesen (2021). https://doi.org/10.34194/geusb.v47.8304 http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 47. 8304. https://doi.org/10.34194/geusb.v47.8304 7 of 8 www.geusbul let in.org deposited. the main glaciodynamic sequence dominating hvideklint relates to an event associated with the advance of the swedish ice stream. we refer to this as the hvideklint glaciodynamic event. however, the presence of two glaciotectonic displaced till units, the ristinge till formation and the klintholm till formation (fig. 6), suggests that an earlier glaciodynamic event affected the area. the old baltic glaciodynamic event, referred to here as the klintholm glaciodynamic event, is responsible for the deposition of these units, where the baltic ice stream transgressed over a glaciolacustrine environment in the hjelm bugt – baltic depression. the only glaciotectonic impact related to this event is the occurrence of a chalk-clay glacitectonite at the base of the till units (fig. 6). an interstadial period of about 10 000 years (houmark-nielsen et al. 2016) between the klintholm and hvideklint glaciodynamic events allowed the large baltic glacial lake to be re-established. at the late glacial maximum, c. 32–25 ka bp, the advancing swedish ice stream caused icebergs to deposit ice-rafted debris in the glaciolacustrine clay (kobbelgård and stubberup have formations) and increased glaciofluvial streams provided sandy deposits, before the ice margin started to deform the hvideklint glaciotectonic complex (pedersen 2000; pedersen & gravesen 2009, 2021). after just a couple of thousand years, the young baltic glaciodynamic event affected the area. the most dramatic effect of this event was the glaciotectonic formation of møns klint itself, and we, therefore, refer to it as the møns klint glaciodynamic event. at hvideklint, an ice lobe in hjelm bugt created the curved shape of the south coast of møn. due to the short-time span between the two glaciodynamic events, only a thin sequence of proglacial sediments was deposited on top of the mid danish till. however, the occurrence of chalk-rich glacitectonite evidently separates this till unit from the lolland till unit, which terminates the glaciodynamic succession at hvideklint. 4 conclusions application of the glaciodynamic sequence stratigraphy allows a systematic reconstruction of the glaciotectonic events that deformed the glacial successions at hvideklint. the two most recent events significantly added to the formation of møn by the early glaciotectonic deformation of the cliffs and the hilly area of west møn, and finally the formation of the hilly area of east møn by the creation of the møns klint glaciotectonic complex. superimposed folding results from the interference of the two events. in this study, we place the top of the chalk (pre quaternary unconformity) to 20–30 m b.s.l. except for the areas, where it is elevated above sea level in the glaciotectonic complexes. the base level for the displaced chalk thrust-sheets is located at the décollement surface, 70–80 m b.s.l. at hvideklint, the chalk is predominantly tectonically brecciated and forms a mixture of chalkglacitectonite and chalk-rich till at the glaciotectonic unconformity truncating the top of the glaciotectonic complex. acknowledgements cand. scient. tove krogh stockmarr is thanked for permission to use parts of her master thesis on hvideklint. dr. dan le heron and an anonymous referee provided helpful suggestions, which improved the text. geodatalog frantz v. platen-hallermund assisted with the arcgis construction of the top-chalk map (fig. 1). funding statement this study was stimulated and supported by the eu horizon2020 project ‘subsol’ (grant agreement no. 642228, www.subsol.org). competing interests none declared. author contributions ssp: writing – original draft; ssp and pg: writing – review and editing, methodology and formal analysis. references aber, j. 1979: kineto-stratigraphy at hvideklint, møn, denmark and its regional significance. bulletin of the geological society of denmark 28, 81–93. houmark-nielsen, m. 1987: pleistocene stratigraphy and glacial history of the central part of denmark. bulletin of the geological society of denmark 36, 1–189. https://doi.org/10.37570/bgsd -198836-01 houmark-nielsen, m. 1994: late pleistocene stratigraphy, glaciation chronology and middle weich selian environmental history from klintholm, møn, denmark. bulletin of the geological society of denmark 41, 181–202. https://doi.org/10.37570/bgsd-1995-41-16 houmark-nielsen, m. 2010: extent, age and dynamics of marine isotope stage 3 glaciation in the southwestern baltic basin. boreas 39, 343–359. https://doi.org/10.1111/j.1502-3885.2009.00136.x houmark-nielsen, m., bennike, o., lamdahl, g. & lüthgens, c. 2016: evidence of ameliorated middle weichselian climate and sub-arctic environment in the western baltic region: coring lake sediments at klintholm, møn, denmark. boreas 45, 347–359. https://doi. org/10.1111/bor.12159 pedersen, s.a.s. 1988: glacitectonite: brecciated sediments and cataclastic sedimentary rocks formed subglacially. in: goldthwait, r.p. & matsch, c.l. (eds): genetic classification of glacigenic deposits, 89–91. rotterdam: a.a. balkema. pedersen, s.a.s. 2000: superimposed deformation in glaciotectonics. bulletin of the geological society of denmark 46, 125–144. https://doi. org/10.37570/bgsd-1999-46-11 pedersen, s.a.s. 2005: structural analysis of the rubjerg knude glaciotectonic complex, vendsyssel, northern denmark. geological survey of denmark and greenland bulletin 8, 192 pp. + 2 plates. https://doi. org/10.34194/geusb.v8.5253 pedersen, s.a.s. 2012: glaciodynamic sequence stratigraphy. in: huuse, m. et al. (eds.): glaciogenic reservoirs and hydrocarbon systems 368, 29–51. london: geological society, special publications. https://doi. org/10.1144/sp368.2 pedersen, s.a.s., fredericia, j. & rasmussen, l.a. 2015: kortbladsbeskrivelse, geologisk kort over danmark, 1:50 000, sakskøbing 1411 https://doi.org/10.34194/geusb.v47.8304 http://www.geusbulletin.org http://www.subsol.org https://doi.org/10.37570/bgsd-1988-36-01 https://doi.org/10.37570/bgsd-1988-36-01 https://doi.org/10.37570/bgsd-1995-41-16 https://doi.org/10.1111/j.1502-3885.2009.00136.x https://doi.org/10.1111/bor.12159 https://doi.org/10.1111/bor.12159 https://doi.org/10.37570/bgsd-1999-46-11 https://doi.org/10.37570/bgsd-1999-46-11 https://doi.org/10.34194/geusb.v8.5253 https://doi.org/10.34194/geusb.v8.5253 https://doi.org/10.1144/sp368.2 https://doi.org/10.1144/sp368.2 pedersen & gravesen 2021: geus bulletin 47. 8304. https://doi.org/10.34194/geusb.v47.8304 8 of 8 www.geusbul let in.org i og 1412 ii syd. geological survey of denmark and greenland map series 6, 42 pp. + kort. pedersen, s.a.s. & gravesen, p. 2006: geologisk kort over danmark 1:50 000, møn. geological map. copenhagen: geological survey of denmark and greenland. pedersen, s.a.s. & gravesen, p. 2009: structural development of maglevandsfald: a key to understanding the glaciotectonic architecture of møns klint, se denmark. geological survey of denmark and greenland bulletin 17, 29–32. https://doi.org/10.34194/geusb.v17.5007 pedersen, s.a.s. & gravesen, p. 2021: kortbladsbeskrivelse, geologisk kort over danmark, 1:50 000, møn dele af 1511 i, 1511 iv og 1512 ii. geus bulletin 48. 8293. (in danish with an english summary). https:// doi.org/10.34194/geusb.v48.8293 pedersen, s.a.s. & petersen, k.s. 1997: djurslands geologi, 96 pp. københavn: danmarks og grønlands geologiske undersøgelse. stockmarr, t.k. 1996: en kvartærgeologisk model over det sydvestlige møn, 164 pp. unpublished msc thesis, university of copenhagen, denmark. (in danish) surlyk, f., rasmussen, s.l., boussaha, m., schiøler, p., schovsbo, n.h., sheldon, e., stemmerik, l. & thibault, n. 2013: upper campanian–maastrichtian holostratigraphy of the eastern danish basin. cretaceous research 46, 232–256. https://doi.org/10.1016/j. cretres.2013.08.006 thomsen, e. 1995: kalk og kridt i den danske underground. in: nielsen, o.b. (red.): danmarks geologi fra kridt til i dag, 31–67. aarhus: aarhus university. (in danish). https://doi.org/10.34194/geusb.v47.8304 http://www.geusbulletin.org https://doi.org/10.34194/geusb.v17.5007 https://doi.org/10.34194/geusb.v48.8293 https://doi.org/10.34194/geusb.v48.8293 https://doi.org/10.1016/j.cretres.2013.08.006 https://doi.org/10.1016/j.cretres.2013.08.006 data article | short bennike & wiberg-larsen 2022: geus bulletin 49. 8294. https://doi.org/10.34194/geusb.v49.8294 1 of 5 a new middle pleistocene interglacial occurrence from ejby, sjælland, denmark ole bennike*1 , peter wiberg-larsen2 1geological survey of denmark and greenland (geus), aarhus, denmark; 2institute for ecoscience, aarhus university, silkeborg, denmark. abstract despite more than a century of investigations, parts of the quaternary stratigraphy of denmark with their fragmented record of deposits remain ambiguous. here we describe a newly found interglacial clay deposit from ejby on sjælland, denmark, from a borehole at 55.695°n, 11.839°e (terrain elevation 5.7 m above sea level). we place the new occurrence on record and provide details of the macrofossil analysis of the sample. the clay contains remains of the present-day temperate bivalve corbicula fluminalis and the caddis fly hydropsyche contubernalis – both inhabiting rivers. the presence of c. fluminalis indicates that the deposit most probably is of middle pleistocene age, older than the last interglacial, the eemian. introduction during the quaternary period that began 2.58 million years, the climate has alternated between glacial and interglacial conditions. interglacial deposits are characterised by remains of plants and animals such as oak and red deer that indicate climatic conditions similar to or warmer than at present, whereas interstadial deposits from glacial stages contain remains of relatively cold-adapted, arctic or subarctic species such as polar willow and reindeer. despite more than a century of investigations, parts of the quaternary stratigraphy of denmark with their fragmented record of deposits remain ambiguous. known interglacial occurrences include the cromerian complex (the harreskovian, marine isotope stage [mis] 19), the holsteinian (mis 11c), the trelde klint interglacial (mis 11), the trianglen interglacial (mis 7), the eemian (mis 5e) and the current interglacial, the holocene (kuneš et al. 2013; odgaard et al. 2016; bennike et al. 2019; candy et al. 2021). moreover, several deposits of unknown age have been described (andersen 1967). in 2018, the geological survey of denmark and greenland (geus) received samples collected at 3-m intervals from a borehole at 55.695°n, 11.839°e (terrain elevation 5.7 m a.s.l.), at the address ejby havnevej 101 on sjælland (fig. 1). shells of the bivalve corbicula were noted in one of the samples; this species indicates a middle pleistocene interglacial age (meijer & preece 2000; bennike et al. 2019), and since such occurrences are rare in denmark, we decided to conduct analyses of macrofossils. this data article puts the new occurrence on record and provides details of the macrofossil analysis of the sample. the identification of fossils from a deposit can help us determine *correspondence: obe@geus.dk received: 08 sept 2021 accepted: 03 jan 2022 published: 02 mar 2022 keywords: middle pleistocene, quaternary, interglacial, river macroinvertebrates, denmark abbreviations: geus: geological survey of denmark and greenland mis: marine isotope stage cal. years bp: calendar years before present a.s.l.: above sea level b.s.l.: below sea level geus bulletin is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: sofie lindström, geus, denmark reviewed by: anne elisabeth bjune (university of bergen, norway), matthew pound (northumbria university, uk). funding: none provided competing interests: see page 4 additional files: see page 4 https://doi.org/10.34194/geusb.v49.8294 https://orcid.org/0000-0002-5486-9946 https://orcid.org/0000-0003-2441-4286 mailto:obe@geus.dk bennike & wiberg-larsen 2022: geus bulletin 49. 8294. https://doi.org/10.34194/geusb.v49.8294 2 of 5 www.geusbul let in.org the age and palaeoenvironment of the deposit. we also briefly review other interglacial and interstadial occurrences from eastern denmark. lithostratigraphy the drilling penetrated clayey till, glaciofluvial gravel, clay with bivalve shells, glaciofluvial gravel and danian limestone (fig. 2). drilling stopped at a depth of 147.5 m below the terrain surface. materials and methods only a single sediment sample was available for analysis of macrofossils from the clay unit with bivalve shells. the sample weighed c. 200 g. it was soaked in a naoh solution at room temperature for 2 days and wet sieved on 0.4-, 0.2and 0.1-mm sieves. the residue left on the sieves was analysed using a dissecting microscope. results and discussion the results of the macrofossil analyses are presented in table 1. the sample was small, and we only found the remains of five taxa. however, some of them are of interest, providing insights into the age and palaeoenvironmental conditions. vascular plants are represented by two taxa. betula sect. albae sp. (tree birch) is common in interglacial assemblages. the small birch nutlets are produced in vast quantities and easily dispersed by wind or rivers. the presence of tree birch indicates a forested environment. zannichellia palustris (horned pondweed) is a small, submerged, perennial water plant that grows in shallow, fresh or brackish waters, in lakes and rivers. it has a wide geographical range, but it does not grow in arctic regions. achenes of z. palustris have previously been reported from interglacial deposits in denmark by hartz (1909) and bennike et al. (2019). invertebrates are represented by shells of the bivalve corbicula fluminalis (fig. 3). the pleistocene occurrence of this species was discussed by meijer & preece (2000), who concluded that the species was present only in temperate stages. although mainly a fresh-water species that lives in rivers, it may have been able to tolerate slightly brackish conditions. it occurred in north-western europe during the early and middle pleistocene, whereas it was absent during the last interglacial. in denmark, it has been found in interglacial deposits from the former free port in copenhagen (johansen 1904), a deposit nowadays referred to mis 7 (bennike et al. 2019). the species has also been reported from an interglacial deposit at førslevgaard in southern sjælland (johansen 1904; location in fig. 1). the age of the deposit at førslevgaard is unknown, but its fauna indicates a middle pleistocene age. fig. 1 map of denmark showing the location of ejby and other middle pleistocene interglacial deposits on sjælland discussed in the text. 10°e 14°e 56°n 57°nsweden denmark germany50 km jylland måløv sjælland fyn førslevgaard trianglen ejby fig. 2 lithological log of the well from ejby (dgu no. 198.964). samples were collected at an interval of 3 m, and the exact thicknesses of the units are unknown. drilling stopped at a depth of 147.5 m below the terrain surface. the star shows the position of the analysed sample. more details on the sediments can be found at https://data.geus.dk/jupiterwww/borerapport. jsp?borid=599264. d ep th b el ow te rra in s ur fa ce (m ) clayey till glaciofluvial gravel glaciofluvial gravel clay with shells danian limestone 0 10 20 30 40 50 table 1 macrofossils from the interglacial deposit at ejby, sjælland. plants betula sect. albae (nutlet) 1 zannichellia palustris (achene) 1 invertebrates daphnia cf. pulex (ephippium) 1 hydropsyche contubernalis (frontoclypeal apotomes) 2 corbicula fluminalis (shells) 10 https://doi.org/10.34194/geusb.v49.8294 http://www.geusbulletin.org https://data.geus.dk/jupiterwww/borerapport.jsp?borid=599264 https://data.geus.dk/jupiterwww/borerapport.jsp?borid=599264 bennike & wiberg-larsen 2022: geus bulletin 49. 8294. https://doi.org/10.34194/geusb.v49.8294 3 of 5 www.geusbul let in.org in addition to bivalve shells, two frontoclypeal apotomes of larvae of the caddisfly (trichoptera) hydropsyche contubernalis were found. like c. fluminalis, h. contubernalis is mainly a fluvial species, nowadays living primarily on the stony or coarse gravel bottom of rivers. however, it is also found in stony, wind-exposed, littoral zones of large lakes, especially in the northern parts of europe (rinne & wiberg-larsen 2017). h. contubernalis nowadays occurs in arresø, the largest lake in denmark (wiberg-larsen 2010). on the undersurface of the substrates mentioned above, larvae build retreats of mineral and organic coarse particles, and in association with these, they spin silky capture nets. the nets function as traps for microand macroinvertebrate prey, transported downstream by the flowing water (or by turbulent currents in lakes). h. contubernalis is nowadays widely distributed throughout europe, from northern scandinavia to spain and the balkans, extending to the british isles in the west and in the east to the urals (neu et al. 2018). there are quaternary danish records from c. 13 300 cal. years bp (coope & böcher 2000) and 10 400 cal. years bp (wiberg-larsen et al. 2001), respectively. in the united kingdom, remains of the species are recorded from late pleistocene deposits (age 35 000−60 000 years bp); langford et al. 2014), and horton et al. (1992) found remains of this and two other hydropsychids in fluvial deposits suggested to be of hoxnian stage ii age. according to ashton et al. (2008), this may correspond to the middle pleistocene, mis 11, corresponding to c. 420 000−370 000 years bp. finally, a single ephippium of daphnia was found. the ephippium is similar to ephippia produced by d. pulex, but it could also represent other species in the daphnia genus. daphnia is a small planktonic crustacean that is often abundant in lakes and ponds, but it can also live in slow-flowing or stagnant parts of rivers. analyses of mitochondrial sequence data (e.g. coi and 16s genes) of aquatic and terrestrial organisms have revealed that the pleistocene was an important period for generating biodiversity, also of species that still exist (e.g. hungerer & kadereit 1998; ribera & vogler 2004; previšic et al. 2009). both c. fluminalis and h. contubernalis are examples of such species. the investigated interglacial clay is a low-energy deposit that likely accumulated in a lake – probably near the mouth of a river as indicated by the two fluvial species. other interglacial and interstadial deposits on sjælland and møn the new investigated site is located c. 500 m south of an overgrown coastal cliff with a layer of stones with an eemian marine fauna that includes paphia senescens (madsen 1968). this last-mentioned deposit is located 10.5 m a.s.l., whereas the new interglacial deposit is located c. 35 m b.s.l., below a succession of clayey till and glaciofluvial gravel. although located in the same area, these deposits appear to belong to two different interglacial stages. other marine faunas referred to the eemian have been reported from nebbegaard in ne sjælland, the former free port in copenhagen, strandegaards dyrehave in se sjælland, stubberup have in eastern møn and hjelm nakke in se møn as well as other sites on møn (fig.  4; madsen et al. 1908; ødum 1933; petersen & konradi 1974; berthelsen et al. 1977; miller & mangerud 1985; pedersen & gravesen 2021). non-marine eemian deposits have also been found on møn (n. hartz in hinze 1937). from slettenshage on røsnæs, nw sjælland, a marine fauna dominated by the marine gastropod turritella communis was described by petersen (1973). based on amino acid analyses, the occurrence was correlated with  holsteinian deposits in nw germany fig. 3 five shells of corbicula fluminalis from ejby. 10 mm https://doi.org/10.34194/geusb.v49.8294 http://www.geusbulletin.org bennike & wiberg-larsen 2022: geus bulletin 49. 8294. https://doi.org/10.34194/geusb.v49.8294 4 of 5 www.geusbul let in.org (miller & mangerud 1985). another interglacial deposit from røsnæs, near skambæks mølle (madsen et al. 1908), may be of the same age. as mentioned earlier, interglacial non-marine deposits referred to mis 7 are known from the former free port in copenhagen and nearby trianglen (bennike et al. 2019; location in fig. 1). other middle pleistocene occurrences have been reported from måløv (bennike et al. 2011; location in fig. 1) and førslevgaard (johansen 1904; hartz 1909). interstadial weichselian non-marine deposits have been reported from kobbelgård and klintholm, se møn (bennike et al. 1994; houmark-nielsen et al. 2016) and sejerø (bennike et al. 2007). marine interstadial deposits with boreoarctic or arctic faunas have been reported from høve (milthers 1900; ødum 1933), holbæk, høng (ødum 1933), nordruplund, bjernede and fjenneslev (ødum 1933; lykke-andersen 1990), gyrstinge (sorgenfrei 1945) and holmstrup (petersen & buch 1974). lists of macrofossils reported from the sites are shown in table s1. conclusions we conclude that the clayey deposit found at ejby was deposited in a low-energy fluvial environment. the mean july temperature was likely higher than in denmark today, and the deposit is clearly interglacial rather than interstadial. the occurrence of the bivalve corbicula fluminalis indicates that the deposit is probably of pre-eemian, middle pleistocene age, at least 200 000 years bp. acknowledgements we thank the reviewers for their valuable corrections to the manuscript. additional information funding statement this study did not receive any funding. author contributions ob: macrofossil analyses, manuscript writing. pwl: identification and interpretation of trichoptera remains as well as supplementary writing/ editing of the manuscript. competing interests none declared. additional files one additional file is available at https://doi.org/10.22008/fk2/qeuaxo references andersen, s.t. 1967: kvartærtiden. istider og mellemistider i danmark. in: nørrevang, a. & meyer, t.j. 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https://doi.org/10.1016/j.quaint.2014.01.037 https://doi.org/10.34194/raekke2.v17.6799 https://doi.org/10.1017/s0016774600021739 https://doi.org/10.1017/s0016774600021739 https://doi.org/10.1016/0277-3791(85)90002-2 https://doi.org/10.1016/0277-3791(85)90002-2 https://doi.org/10.34194/geusb.v35.4925 https://doi.org/10.34194/geusb.v48.8293 https://doi.org/10.1111/j.1365-294x.2008.04046.x https://doi.org/10.1111/j.1365-294x.2008.04046.x https://doi.org/10.1046/j.1365-294x.2003.02035.x https://doi.org/10.1111/j.1502-3885.2001.tb01049.x https://doi.org/10.1111/j.1502-3885.2001.tb01049.x https://doi.org/10.34194/raekke4.v2.6980 https://doi.org/10.34194/raekke4.v2.6980 research article | short svennevig et al. 2020: geus bulletin 44. 5302. https://doi.org/10.34194/geusb.v44.5302 1 of 6 preliminary landslide mapping in denmark indicates an underestimated geohazard kristian svennevig*1, gregor luetzenburg2, marie k. keiding1, stig asbjørn schack pedersen1 1geological survey of denmark and greenland (geus), copenhagen, denmark. 2department of geosciences and natural resource management, university of copenhagen, copenhagen, denmark abstract the process of coastal erosion is well known to the public and decision-makers in denmark; however, there is little awareness of the risks posed by larger landslides. only a few scientific studies investigate landslides in denmark, and as a result, the country is underrepresented in international landslide inventories. here, we present a systematically produced preliminary landslide inventory based on digital elevation models and high-resolution orthophotos. so far, the preliminary inventory documents 3026 morphological expressions of landslides close to the coast and inland, showing that landslides are more widespread in denmark than previously recognised. a number of these landslides are near buildings and infrastructure. this paper therefore highlights the potential for geohazardous landslides to occur in denmark on a national scale and discusses some of the implications. two of the major questions arising from this study are (1) how to approach potential geohazards in a country with no framework or precedence for landslide hazard and risk management and (2) how landslides and associated risk in denmark will evolve under a changing climate. introduction previous studies of landslides in denmark are limited. most are mainly field-based investigations of single events or areas with a focus on rockfalls (table 1), and most of the sparse information is published in danish ‘popular science’ magazines. as a result, denmark may be underrepresented in international landslide inventories (herrera et al. 2018; mateos et al. 2020). based on the data and methods available at the time, pedersen et al. (1989) suggested that landslides are not a serious problem in denmark. nadim et al. (2008) indicated landslides as the main geohazard in denmark, but that nonetheless, the country was not seriously affected by them. these studies laid important groundwork by identifying areas prone to slope failure and provided the first insights into process behaviour at selected sites. nevertheless, their findings also raised the question: ‘how abundant are landslides in denmark, actually?’ with the emergence of nationwide digital elevation models (dems), we can start to acquire comprehensive knowledge about surface processes all over *correspondence: ksv@geus.dk received: 18 june 2020 accepted: 15 sept 2020 published: 09 nov 2020 keywords: climate change, denmark, geohazards, landslide inventory, landslides abbreviations: dem: digital elevation model gis: geographical information system dod: dem of difference insar: interferometric synthetic aperture radar geus bulletin is an open access, peerreviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: rasmus bødker madsen (geus, denmark) reviewed by: kurt h. kjær (university of copenhagen, denmark) and reginald l. hermanns (geological survey of norway, norway) funding: see page 6 competing interests: none declared additional files: none provided https://doi.org/10.34194/geusb.v44.5302 mailto:ksv@geus.dk svennevig et al. 2020: geus bulletin 44. 5302. https://doi.org/10.34194/geusb.v44.5302 2 of 6 www.geusbul let in.org the country. here, we align danish landslide research with international landslide research by applying stateof-the-art landslide classifications and methods for landslide inventory mapping. this work is the first step towards a comprehensive mapping of landslides in denmark. we consider the current landslide mapping to be preliminary, because it is not yet validated and extended with additional information about the mapped slides. in the future, the database will be validated, extended and published in its entirety to address the various scientific questions raised at the end of this paper. methods the current landslide mapping is based entirely on freely available high-resolution geodata from the danish agency for data supply and efficiency (sdfe; table  2). the primary dataset used is a 40 cm spatial resolution dem produced from airborne laser scanning data (sdfe 2020). these data have been acquired since 2007 and are periodically updated. we used the 2015 dem as it was the latest complete release available when our project began. the dem is visualised as a multidirectional hillshade model in order to observe morphological appearances of landslides across denmark. orthophotos with a 12.5 cm resolution from multiple years assisted the mapping for visual validation. generally, spatial coverage is consistent throughout denmark. in a few small areas, for example, steep or overhanging cliffs on stevns klint and møns klint, data accuracy is insufficient due to the close to nadir angles of the airborne sensors. in other areas, the model suffers from insufficient data points causing morphologies to be obscured by interpolation. overall, these minor data gaps do not influence the spatial completeness of the mapped landslide database. in the few cases where landslides have been mapped in these areas, it has been noted in the attribute table. the mapping workflow is similar to that reported by svennevig (2019) for mapping landslides in greenland and a simplified version to that of slaughter et al. (2017). a landslide was mapped when either a scarp or a displaced unit or both were clearly visible in the multidirectional hillshade model (fig. 2a). we thus only include landslides with moderate to high confidence, similar to burns & madin (2009) and slaughter et al. (2017). moreover, based on the high-quality and detail of the dem, as well as the support of the observations with the time series of orthophotos (table 2), we consider all the mapped landslides to be identified with high confidence. we characterised each landslide by its morphological features following the idealised depiction and table 1 previously published peer-reviewed studies on landslides in denmark. see fig. 1 for locations reference area topic andersen (1957) salten geomorphology hansen (1959) salten geomorphology prior (1973) røsnæs landslides, mudslides hansen (1975) salten geomorphology prior & eve (1975) røsnæs landslides, mudslides prior (1977) røsnæs, helgenæs and røjle klint mudslides pedersen (1987) mors glaciotectonic complex, landslides pedersen et al. (1989) regional landslides hutchinson (2002) møn rockfalls busby et al. (2002) møn, stevns rockfalls pedersen & møller (2004) møn, stevns rockfalls pedersen (2005) lønstrup klint glaciotectonic complex, landslides nadim et al. (2008) regional geohazards pedersen & gravesen (2009) møn rockfalls pedersen & damholt (2012) stevns rockfalls pedersen (2012) møn rockfalls table 2 publicly available data from the danish agency for data supply and efficiency used in the landslide mapping name type publication year access source resolution (cm) geodkorto2019 orthophoto 2019 wms cowi 12.5 geodkorto2018 orthophoto 2018 wms cowi 12.5 geodkorto2017 orthophoto 2017 wms cowi 12.5 geodkorto2016 orthophoto 2016 wms cowi 12.5 geodkorto2015 orthophoto 2015 wms cowi 12.5 denmark’s elevation model dem (2007) 2015 wcs sdfe (160) 40 ddoland2014 orthophoto 2014 wms cowi 12 https://doi.org/10.34194/geusb.v44.5302 http://www.geusbulletin.org svennevig et al. 2020: geus bulletin 44. 5302. https://doi.org/10.34194/geusb.v44.5302 3 of 6 www.geusbul let in.org nomenclature of a rotational earthflow from highland (2004), as well as the commonly used landslide classification of hungr et al. (2014). in many cases, additional morphological features such as a crown, transverse cracks, main body or foot supported the identification. when the foot was not visible in the hillshade model due to erosion by waves or anthropogenic overprint, the older orthophotos were consulted to identify the landslides morphological expression on the record closest to the event. it should be noted that as the mapping is based on the above described morphological criteria and datasets, it is not suitable for mapping slides with small volumes or faint morphologies such as rockfalls and mudslides. these types are thus expected to be underrepresented in the database. the mapping itself consists of drawing a polygon with a high number of vertices around an identified landslide, based on the spatial delineation in the high-resolution hillshade model. landslides with a minimum spatial extend of around 25 m2 were mapped. two mappers each mapped around half of the landslides in the preliminary database and validated the respective other half. we used qgis, a free and open source cross-platform geographic information system (gis) to display the multidirectional hillshade model and the orthophotos, as well as to map the landslides. every mapped polygon of a landslide was assigned a unique identification number, and the following information was recorded for each id: shape, proximity to coastline, morphological indication of recent activity, given name (if available), name of the person who mapped the landslide and hazard potential. in a free-text column, we noted additional useful information, including anthropogenic cause/ overprint, type of hazard (e.g. ‘house on slide’), validation by historic sources, sign of change in the dataset used and dem quality. preliminary results and discussion at the time of publication, the preliminary landslide inventory consists of 3026 distinct landslides in denmark (fig. 1). of these, 2318 are at or near the coast (<300 m) and 708 are more than 300 m from the coast and considered inland. 56° 56° 12° 9° 12° 9° 50 km 55° 15° n jylland fyn mariager fjord lønstrup klint røsnæs københavn røjle klint aabenraa helgenæs mors møns klint stevns klint sjælland bornholm fig. 2a fig. 2c fig. 2d fig. 2b salten vejle fig. 1 mapped landslides in denmark. green dots show 172 landforms with overprint of younger processes, such as fluvial incision. red dots show 2854 landslides that are not modified by other geological processes – some of these show indications of recent activity. place names mentioned in the text and table 1 along with positions of panels in fig. 2 are shown. https://doi.org/10.34194/geusb.v44.5302 http://www.geusbulletin.org svennevig et al. 2020: geus bulletin 44. 5302. https://doi.org/10.34194/geusb.v44.5302 4 of 6 www.geusbul let in.org fig. 2 four examples of landslides from denmark in the dem hillshade (left) and an orthophoto (right). a: large active rotational landslide (r) and flow (f), east of røjle klint, fyn. b: large rotational landslide, north of mors, northern jylland; the southern part of the landslide is partially concealed by farming activities. the northern uncultivated part of the landslide has clear internal structures, while the southern part is ploughed annually obscuring structures. this landslide would not have been mapped if not for the structures in the northern part. c: inland landslide near vejle, jylland, superimposed by smaller flows. the large landslide is presumably a rotational or translational landslide or a combination. d: example of a landslide (green) eroded by fluvial incision and a younger landslide (red) in mariager fjord, jylland. structures in the young landslide are somewhat obscured by quarrying (q), which could have been a triggering factor for the landslide. 200 m 200 m 200 m 200 m a b c d r f q n https://doi.org/10.34194/geusb.v44.5302 http://www.geusbulletin.org svennevig et al. 2020: geus bulletin 44. 5302. https://doi.org/10.34194/geusb.v44.5302 5 of 6 www.geusbul let in.org many of the mapped landslides may record multiple events that overlap in extent. as such, the mapped polygons should be considered areas of past and/or present landslide activity. some of these may have been successive events, while others are single events of abrupt change. others may represent areas of continuous slow sliding. anthropogenic overprint, such as from cultivation, buildings and infrastructure, is commonly observed on many landslides and can obscure their expression in the dem. since farmland covers 56% of denmark’s surface, the chance that landslide structures are obscured by cultivation and thus not mapped is rather high (fig. 2b). this is especially the case for shallow slides and inland landslides (fig. 2c). quarries can be overprinted by slides, potentially having a triggering effect (fig. 2d), but old unmapped quarries may also be mistaken for landslides. a total of 172 landforms, some exceptionally large, fit the morphological criteria for a landslide but have a morphological overprint by younger processes such as fluvial incision (see fig. 1). these may be interpreted as landslides formed during periglacial or paraglacial phases of landscape development (fig. 2d). these extraordinary landforms resemble large-scale landforms in aabenraa, southern denmark, identified as possible landslides by lykke-andersen & sørensen (2018) and could have played a hitherto unrecognised role in degrading the danish landscape in postglacial times, which calls for further investigation. outlook since this paper is the first step towards a comprehensive landslide inventory of denmark, there are several means to extend the work in terms of verification, expansion and application of the database. verification and expansion of the database the preliminary landslide inventory will be quality controlled for all of denmark and validated in places by means of field observations and comparison with mentions of landslide events in newspapers and historical archives. linking hydrological, meteorological, land use and geological datasets to the landslide inventory could provide further information about the preconditioning and driving factors of slope failures in denmark and elsewhere. this could lead to the first landslide susceptibility map of denmark. the current database does not provide any information about the state of activity of the landslides. a valuable add-on will thus be a classification of each landslide as either active, suspended, reactivated, dormant, stabilised or relict sensu cruden & varnes (1996), where possible. this will be achieved by the analysis of interferometric synthetic aperture radar (insar), dem of difference (dod) and multiple generations of orthophotos (table 2). sentinel 1 insar data from the european copernicus satellite mission will be used to calculate the line of sight displacement of slides with natural reflectors on a 6-day repeat cycle since 2015, enabling an unprecedented temporal resolution for the monitoring of current surface movements. dods will be produced from the three nationwide dems recorded at 4–8-year intervals since 2006. they can also be used to quantify three-dimensional morphological changes and rates of erosion. the dod coverage can be extended up to 90 years back in time on a local scale by implementing dems from several sets of historical aerial stereo photos, providing valuable information on the longer-term evolution of landslide activity in denmark. application the high temporal and spatial resolution of the underlying datasets available for the mapping of landslides in denmark has several potential applications. landslide activities across the globe are expected to increase in the coming decades as a result of climate change (ipcc 2014; gariano & guzzetti 2016). in denmark, climate change is expected to cause an increase in frequency and magnitude of heavy precipitation events, storm surges and sea-level rise, which may have an accelerating effect on landslide activity. it is crucial to understand how these processes affect landslides and develop a toolset to predict future changes and quantify how these will impact society. combining local and regional climate models with the above-mentioned susceptibility map would allow us to quantify the effects of climate change on landslides in denmark now and in the near future. the database could also be applied to quantify the threat landslides and may pose to danish society. during the mapping exercise, several infrastructure objects, buildings and agricultural lands were noted as located on or near landslides. integrating dod and insar data into the database, as mentioned earlier, along with overlay analysis of gis datasets of infrastructure, buildings and land use will help to identify potentially hazardous landslides. these data analyses should be followed up by detailed site-specific field work to form the basis for informed decisions on mitigation measures. this work can further lay the foundation for the development of a legislative framework dealing with the consequences of landslides in denmark and integrate a landslide hazard and risk mitigation evaluation in climate adaptation plans. this is an area, in which denmark has previously been identified as falling behind other eu countries (mateos et al. 2020). https://doi.org/10.34194/geusb.v44.5302 http://www.geusbulletin.org svennevig et al. 2020: geus bulletin 44. 5302. https://doi.org/10.34194/geusb.v44.5302 6 of 6 www.geusbul let in.org acknowledgements we thank professor aart kroon and assistant professor anders a. bjørk for their valuable input on the project. sdfe are thanked for their support and joanna balasis-levinsen (sdfe) for giving us early access to insar data. we kindly thank professors kurt h. kjær and reginald hermanns for constructive reviews that helped to improve the final version of the paper. additional information funding statement gl has received funding from the european union’s horizon 2020 research and innovation programme under the marie skłodowska-curie grant agreement no 801199, which funded his hours on the study. authors’ contributions ksv: conceptualization. ksv, gl: data curation, formal analysis and writing (original draft, review and editing). mkk: formal analysis and writing (review and editing). sasp: writing (original draft). references andersen, s.a. 1957: de jyske kildedale og deres problemer. meddelelser dansk geologisk forening 13, 438–440. burns, w.j. & madin, i.p. 2009: protocol for inventory mapping of landslide deposits from light detection and ranging (lidar) imagery. oregon department of geology and mineral industries special paper 42, 30 p. busby, j.p., gourry, j.c., senfaute, g., pedersen, s.a.s. & mortimore, r. 2002: can we predict coastal cliff failure with remote, indirect measurements? in: jakeways, j. & mcinnes, r. 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(eds): landslides and engineered slopes: protecting society through improved understanding, 1127–1132. london: taylor & francis group. pedersen, s.a.s. & damholt, t. 2012: cliff collapse at stevns klint, southeast denmark. geological survey of denmark and greenland bulletin 26, 33–36. https://doi.org/10.34194/geusb.v26.4745 pedersen, s.a.s., foged, n. & frederiksen, j. 1989: extent and economic significance of landslides in denmark, faroe islands and greenland. in: brabb, h. (ed): landslides: extent and economic significance, 153–156. rotterdam: balkema. pedersen, s.a.s. & gravesen, p. 2009: structural development of maglevandsfald: a key to understanding the glaciotectonic architecture of møns klint, se denmark. geological survey of denmark and greenland bulletin 17, 29–32. https://doi.org/10.34194/geusb.v17.5007 pedersen, s.a.s. & møller, i. 2004: prediction and risk evaluation of chalk cliff collapse: the protect project. geological survey of denmark and greenland bulletin 4, 89–92. https://doi.org/10.34194/geusb.v4.4793 prior, d.b. 1973: coastal landslides and swelling clays at røsnæs, denmark. danish journal of geography 72, 45–48. https://doi. org/10.1080/ 00167223. 1973.10649024 prior, d.b. 1977: coastal mudslide morphology and processes on eocene clays in denmark. danish journal of geography 76, 14–33. https://doi. org/10.1080/00167223.1977.10649071 prior, d.b. & eve, r.b. 1975: coastal landslide morphology at røsnæs, denmark. danish journal of 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https://doi.org/10.1007/s10346-013-0436-y https://doi.org/10.1130/reg15-p257 https://doi.org/10.1130/reg15-p257 https://doi.org/10.1016/j.landurbplan.2019.103740 https://doi.org/10.18814/epiiugs/2008/v31i1/024 https://doi.org/10.1144/gsl.sp.1987.029.01.14 https://doi.org/10.1144/gsl.sp.1987.029.01.14 https://doi.org/10.34194/geusb.v8.4848 https://doi.org/10.34194/geusb.v8.4848 https://doi.org/10.34194/geusb.v26.4745 https://doi.org/10.34194/geusb.v17.5007 https://doi.org/10.34194/geusb.v4.4793 https://doi.org/10.1080/​00167223.​1973.10649024 https://doi.org/10.1080/​00167223.​1973.10649024 https://doi.org/10.1080/00167223.1977.10649071 https://doi.org/10.1080/00167223.1977.10649071 https://doi.org/10.1080/​00167223.1975.10649043 https://doi.org/10.1080/​00167223.1975.10649043 https://sdfe.dk/hent-data/danmarks-hoejdemodel/ https://doi.org/10.34194/geusb-201943-02-07 preliminary landslide mapping in denmark indicates an underestimated geohazard abstract introduction methods preliminary results and discussion outlook verification and expansion of the database application acknowledgements additional information references tables table 1 previously published peer-reviewed studies on landslides in denmark. see fig. 1 for location table 2 publicly available data from the danish agency for data supply and efficiency used in the l figures fig. 1 mapped landslides in denmark fig. 2 four examples of landslides from denmark in the dem hillshade (left) and an orthophoto (right geological survey of denmark and greenland bulletin 35, 2016, 35-38 35© 2016 geus. geological survey of denmark and greenland bulletin 35, 35–38. open access: www.geus.dk/publications/bull roskilde fjord is a characteristic n–s-trending geomorphological element in north-east sjælland (fig. 1). the eastern coastline of the fjord forms a nearly straight, sse– nnw-trending lineament from the town of roskilde to the coastal areas at kattegat. due to the records from wells, it has long been known that a fault zone has to be present along this lineament (bondesen 1979). the fault is named after the fjord: the roskilde fjord fault (fig. 2). however, a detailed analysis of the well data in the roskilde area has shown that a number of minor faults are present, superposed by various landscape elements. these are the inner fjord and valley distributary at roskilde, the kornerup å valley, the inlet of lejre vig and its contributory, and the peninsula of bognæs (figs 1, 2). in this paper we propose a lithostratigraphic division of the cenozoic deposits and their dynamic development in the roskilde area. a distribution of the faults with estimated maximum displacements is presented, and their relations to the geomorphological features are outlined. the description is based on several years’ studies of the roskilde fjord fault complex and presented here due to the recent interest in neotectonics in denmark and environmental considerations focused on roskilde fjord (gravesen & pedersen 2005). stratigraphy of the roskilde area the palaeogene and quaternary stratigraphy of the area is briefly outlined here based on well data and a few sand and gravel pit outcrops (gravesen & pedersen 2005). the palaeogene deposits in the area include danian and selandian formations (65.5 to 58.7 ma). the late danian københavn kalk formation consists of calcisiltitic and calcarenitic yellow-white limestone with some thick chert layers (thomsen 1995). the overlying, early selandian deposits were laid down after a period of erosion and often begin with glauconitic conglomerate and greensand of the lellinge greensand formation followed by olive grey glauconitic silty limestone, sand and clay (clemmensen & thomsen 2005). the pre-quaternary unconformity thus truncates danian as well as selandian units. the oldest quaternary deposits are probably from the late saalian (150–130 ka bp). the clayey tills and sand layers are found in deep buried valleys down to 80 m below the surface. the deposits are known from wells but are difficult to date precisely. jacobsen (1985) also suggested that saalian deposits were the oldest quaternary deposits in the area. from the early middle weichselian thin and fragmented sand layers were deposited in rivers and lakes, and the surficial processes possibly initiated the valley formation c. 70–55 ka bp. a middle weichselian ice stream from the east, the 55–50 ka bp old baltic ice stream, deposited the ristinge klint till formation, characterised by reddish staining of the clayey and sandy tills (houmark-nielsen 2010). the late weichselian himmelev formation consists of cross-bedded meltwater sand and gravel deposited by braided streams from the north (jacobsen 1985), indicattectonic control on the formation of roskilde fjord, central sjælland, denmark stig a. schack pedersen and peter gravesen bognæs eskilsø sjælland kattegat le jre v ig ko rn er up å himmelev kirke såby veddelev roskilde risø fig. 1. terrain model of the roskilde area and index map of denmark showing the location of the area in focus. note the strong s–n-trending lineament including the linear coast lines of roskilde fjord. 3636 ing deposition at the southern front of the melting kattegat ice stream at c. 26 ka bp. this unit is overlain by the grey to brown, massive mid danish till formation deposited by the ice stream from north-east during the last glacial maximum (houmark-nielsen 1987, 2010). the mid danish till formation is overlain by the hedeland formation, which is a very coarse-grained, cross-bedded sand/gravel formation deposited as a proximal sandur and in meltwater channels (jacobsen 1985). the last ice advance, the young baltic ice advance, was separated into two. the first, the east jylland advance from the south-east and east deposited the east jylland till formation of grey and brown clayey till at 19–18 ka bp (houmark-nielsen 2010). in the following period, which was free of ice cover, glaciolacustrine sediments were deposited. the last advance from the east and south-east deposited the brown, clayey bælthav till formation (18–17 ka bp). during the final melting of the last glaciers the vindinge formation consisting of fine-grained sand, silt and clay was deposited in a dead ice landscape. marine holocene sand and mud deposits are found along the coast of roskilde fjord, while freshwater deposits occur along small streams and in lakes. method of cross-section construction the analysis of the subsurface geology in the roskilde fjord area was carried out by the construction of a large number of mainly e–w-trending cross-sections perpendicular to the interpreted strike of fault features. the cross-sections were constructed in arcgis©, using well data from the jupiter database along chosen sections with a bandwidth of 1 km. maps were constructed from a selection of these wells, providing information on the depth to the pre-quaternary unconformity and the local lithology at the pre-quaternary surface (fig. 2). the pre-quaternary unconfomity map was produced by an integration over the well data points using arcgis©. depth to the pre-quaternary unconformity the map of the depth to the pre-quaternary unconformity (fig. 2) shows that the surface forms a high plateau to the east of roskilde, reaching a level above 30 m a.s.l. along the drainage creek to the south-eastern arm of roskilde fjord and along the eastern coastline, the structural contours of the surface of the pre-quaternary unconformity are closely spaced and trend n–s, indicating an escarpment parallel to the east coast of the roskilde fjord. below the central part of the southern depression of roskilde fjord, where the peninsula of bognæs is located, the depth to the pre-quaternary unconformity is more than 40 m b.s.l. thus, the level change from east to the central part of roskilde fjord is almost 80 m. the south-western corner of the roskilde fjord is dominated by two strong lineaments parallel to the geomorphic features kornerup å and lejre vig, where two buried val0 -30 -20 -20 -20 -20 -10 -40 10 10 -5 0 20 40 -40 0 -10 -4 0 -50 -20 0 -10 30 0 -10 0 -10 -6 0 -6 0 -70 -30 -30 -2 0-3 0 -2 0 -20 -20 -10 -20 30 -10 -3 0 -10 10 -40 -40 ri sø f au lt h or ns he rr ed f au lt le jre v ig f au lt g ev ni ng e f au lt r oskilde fjord fault h erslev fault kattinge vig graben søndersø graben søndersø graben la ng ho lm h or st ko rn er up g ra be n ågerup block roskilde fjo rd 5 kmn normal fault lineament paleocene deposits danian limestone w ew e himmelev block fig. 2. geological map of the formations occurring on the pre-quaternary unconformity surface and location of the cross-section shown in fig. 4. the contour lines show the depth in metres to the pre-quaternary unconformity based on well data. roskilde forkastning risø fault ågerup block himmelev block veddelev block c. 100 m c. 1 km roskilde fjord fault fig. 3. the block diagram illustrates the relay faults along the east side of roskilde fjord. 37 leys are eroded to a depth of more than 70 m b.s.l. (fig. 1). between the lineaments the pre-quaternary unconformity has an elevation of 0–10 m b.s.l., and farther to the west and south-west at kirke såby the surface forms a plateau at about sea level (fig. 2). to the north the roskilde fjord changes from a broad basin to a narrow fjord arm extending northwards. the shift is located to the area around eskilsø, and the broad basin depression to the south is interpreted as the interference between the roskild fjord lineaments and the e–w-trending søndersø graben (fig. 2). this graben system is interpreted as partly tectonically controlled but mainly mirrors a tunnel-valley drainage system to the west related to the baltic ice advance. block faulting of danian limestone and paleocene limestone and clay only two formations are distinguished at the pre-quaternary unconformity surface: danian limestone and paleocene limestone and clay (fig. 2). their distribution follows two rules: danian limestone either appears where a trench has been carved down to a depth below the base of the paleocene limestone, greensand and clay, or it appears on elevated plateaux above the paleocene deposits. the latter is the case east of roskilde, whereas carved trenches are seen below the kornerup å valley and the lejre vig buried valley (figs 1, 2). the broad, e–w-trending søndersø graben belongs in the first group representing a buried valley sandur. the faults are relay faults (korstgård 1996). this means that they have a hinge point where the displacement is negligible, and from where the displacement increases along strike (fig. 3). a culmination of the displacement occurs about 1–4 km from the hinge point, and then the displacement decreases to another hinge point at the opposite termination of the fault. in the roskilde fault complex the maximum displacement is calculated to about 60 m, but most faults have displacements only in the range of 10–30 m. the marker horizon used for estimation of the displacement is the boundary between the danian limestone and the selandian deposits. this boundary is believed to have been an almost horizontal plane, given that both units are marine successions deposited in the broad shelf environment that dominated the danish basin in the palaeogene. w e gl. lejre kornerup å glim vor frue 2 km 0 -10 -20 -30 -40 -50 m bælthavs till fm weichselian saalian paleocene east jutland till fm hedeland fm mid danish till fm himmelev till fm ristinge till fm undifferentiated glaciofluvial deposits undifferentiated glaciofluvial till fm selandian limestone, greensand and clay unconformity danian limestone fault fig. 4. e–w cross-section of the area south of roskilde fjord (location shown in fig. 2). below the geological cross-section, two back-stripped cross-sections illustrate the early fault and landscape evolution. grey vertical lines represent wells. 3838 glaciodynamic development of the quaternary succession the interpretation of the dynamic development includes a palinspastic reconstruction shown in fig. 4. the interpreted cross-section also illustrates two former structural steps in the development. the first of these goes back to an unconformity that is interpreted as a peneplanation in the main part of the middle-late weichselian. the second step is a further back-stripping to the supposed eemian peneplanation after the saalian glaciation; this scenario illustrates the interpreted faulting of the boundary between the danian limestone and selandian unit. evidently, at this time the roskilde fault complex had already been active. furthermore, a deep incised valley along the kornerup å lineament was probably initiated in the neogene prior to the late pleistocene. we suggest that deep valley erosion in the elsterian was succeeded by infill into the valley of meltwater sand and tilly material in the saalian (fig. 4). during the middle weichselian the ristinge advance deposited meltwater sand and the ristinge till formation on a succession that had been faulted. in the beginning of the late weichselian, the last glacial maximum initiated the swedish ice advance (pedersen 2012), which resulted in meltwater streams eroding into the surfaces of downthrown blocks. renewed erosion and deposition took place in the kornerup å valley, and new valleys were formed in the blocks to the east and to the west along fault lineaments. after the meltwater sand of the himmelev formation had filled the valleys to form a relatively even outwash plane surface, the mid danish till formation covered the area with a more than 10 m thick clayey basal till. after the deposition of the mid danish till formation, a new phase of faulting occurred along the roskilde fault. this resulted in a relative uplift of the block to the east and the total vertical displacement by c. 60 m. during the young baltic ice advance in the late weichselian the mid danish till formation and the underlying himmelev formation were eroded away from the elevated area to the east, whereas the sequence was preserved in a down-faulted block to the west. the young baltic ice advance had oscillating advances and retreats including the east jutland ice advance and the bælthav ice advance (houmark-nielsen 2010). at this time the main part of eastern denmark became covered by a huge outwash plane. at roskilde this resulted in the deposition of the hedeland formation which is the main source for gravel exploitation in the roskilde area (the previous interpretation of the hedeland formation by jacobsen 1985 is now considered redundant). thin beds of till are commonly present in this meltwater succession, representing the oscillating ice conditions. before the final erosion of the valleys and fjord arms the bælthav till formation was deposited shaping the present hummocky moraine plateau landscape. conclusions the geomorphology of the roskilde area mirrors a number of tectonic and depositional elements in the subsurface. the fault zones responsible for several terrain lineaments are relay faults that outline an imperfect en échelon pattern. the fault displacements can be calculated from the position of the boundary between the danian limestone and the selandian unit. neotectonic displacement was active shortly after the deposition of the mid danish till formation. the deepest buried valleys were probably eroded out during the elsterian glaciation and filled with sand and till during the saalian glaciation. these valleys were buried during the weichselian glaciodynamics. references bondesen, e. 1979: roskilde, by og landskab – geologi og samfund. in: birkebæk, f.a. (ed.): 13 bidrag til roskilde byog egnshistorie. roskilde museums 50 års jubilæumsskrift, 20–41. roskilde: roskilde museum. clemmensen, a. & thomsen, e. 2005: palaeoenvironmental changes across the danian–selandian boundary in the north sea basin. palaeogeography, palaeoclimatology, palaeoecology 219, 351–394. gravesen, p. & pedersen, s.a.s. 2005: de geologiske forhold ved risø. redegørelse udarbejdet på basis af eksisterende data. danmarks og grønlands geologiske undersøgelse rapport 2005/30, 40 pp. houmark-nielsen, m. 1987: pleistocene stratigraphy and glacial history of the central part of denmark. bulletin geological society of denmark, 36, 189 pp. houmark-nielsen, m. 2010: extent, age and dynamics of marine isotope stage 3 glaciations in the southwestern baltic basin. boreas 39, 343–359. jacobsen, e.m. 1985: en råstofgeologisk kortlægning omkring roskilde. dansk geologisk forening, årsskrift for 1984, 65–78. korstgård, j.a. 1996: ekstensionsforkastninger. geologisk tidsskrift 1, 1–24. pedersen, s.a.s. 2012: glaciodynamic sequence stratigraphy. in: huuse, m. et al. (eds): glaciogenic reservoirs and hydrocarbon systems. geological society, london, special publications 368, 29–51. thomsen, e. 1995: kalk og kridt i den danske undergrund. in: nielsen, o.b. (ed.): danmarks geologi fra kridt til i dag. aarhus universitet geokompendier 1, 32–67. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: sasp@geus.dk geological survey of denmark and greenland bulletin 15, 2008, 13-16 a high resolution carbon isotope (δ13c) profile through the upper campanian to maastrichtian chalk was recently completed based on material from the stevns-1 core from the stevns peninsula, eastern denmark. the δ13c variation of marine carbonates essentially reflects global perturbations in the carbon cycle, i.e. the burial fluxes of carbonate carbon versus organic carbon. it is widely observed that the δ13c variation broadly tracks the eustatic sea-level curve, and that δ13c curves can be used for stratigraphic correlation (e.g. jarvis et al. 2002). in the stevns-1 core, a total of 29 notable isotope changes have been identified in the upper cam panian to maastrichtian succession. in order to evaluate the stratigraphic significance of the isotope changes, the variation in δ13c values of the mid-maastrichtian chalk from cores in eastern denmark and the danish north sea, and from outcrops at rørdal, northern jylland has been examined (fig. 1). the selected interval is characterised by distinct chalk and marl cycles in the stevns-1 and karlslunde-1 cores and in the rørdal quarry (fig. 2), whereas a non-cyclic clean chalk is found in the m-10x well from the north sea. in the rørdal quarry, the chalk–marl unit spans the upper–lower maa strichtian boundary in the boreal brachiopod and belemnite stratigraphies (surlyk 1984; unpublished data, b. lauridsen & f. surlyk). in stevns-1 and karlslunde-1 the chalk–marl unit was deposited during the younger part of nannofossil subzone uc20b (sheldon 2006, in press). this paper presents preliminary results of a high-resolution study of carbon isotopes, carried out by the geological survey of denmark and greenland (geus) in co-operation with partners from the department of geography and geology at the university of copenhagen. this paper is a product of the cretaceous research centre (crc) at geo center denmark. © geus, 2008. geological survey of denmark and greenland bulletin 15, 13–16. available at: www.geus.dk/publications/bull correlation of carbon isotope events in the danish upper cretaceous chalk niels h. schovsbo, susanne lil rasmussen, emma sheldon and lars stemmerik 13 50 km n <500 m 500–1000 m 1000–1500 m 1500–2000 m >2000 m thickness of chalk –gnibøkg nir nisa bhsina d ros neg nrot–ier f enoztsi uq outer limit fault salt dome late cretaceous inversion zone basement high nyf hgih a r s k copenhagen jylland 4°00' 5°00' 55°30' 56°00' dan fieldm-10x 25 km fault zone inversion zone oilfield in chalk gasfield in chalk studied well uk norway denmark germany the netherlands 500 km b fig. 1. a: thickness of the upper cretaceous to danian carbonates in the danish area. k, karlslunde-1 borehole; s, stevns-1 borehole; r, rørdal quarry. modified from stemmerik et al. (2006). b: location of the m-10x well and the dominant late cretaceous structural elements in the danish central graben (modified from ineson et al. 2006). the inset shows the territorial borders in the north sea. stevns-1 carbon isotope curve the stevns-1 borehole represents the first continuously cored maastrichtian section in north-western europe (stemmerik et al. 2006). the borehole is located close to the stevns klint coastal cliff section (fig. 1) and penetrated 456.1 m of lower danian bryozoan limestone and uppermost maastrichtian to campanian chalk with 100% core recovery. the campanian to maastrichtian chalk is composed of very pure carbonate (90–95%) and consists mainly of coccoliths. marl beds (1–10 cm thick) occur in two intervals, in the campanian to maastrichtian boundary interval and within the midmaastrichtian. the δ13c values in the stevns-1 core have been measured on bulk sediment samples collected with a density of approximately four samples per metre. isotope measurements were made at the department of geography and geology. the analytical precision of the δ13c values is better than ±0.05‰. all carbon isotope results are given relative to the v-pdb standard. the mid-maastrichtian marl succession in stevns-1, from 90–110 m, forms part of a 120 m thick interval characterised by high δ13c values that range between 2.2 and 2.5‰ (fig. 3). the interval is characterised by a gentle upward decrease in δ13c values terminated by a marked fall at 85 m, near the boundary between the nannofossil subzones uc20b and uc20c. the overlying chalk, referred to subzone uc20c, is characterised by lower and more variable δ13c values between 1.7 and 2.2‰, and shows a general fall towards a distinct low at the uc20c–d boundary (fig. 3). within the interval from 135 m to 60 m in stevns-1, we have defined seven conspicuous isotope events in chalk belonging to subzones uc20b and uc20c (fig. 3). the three lower events, uc20b1–uc20b3, all predate the onset of marl deposition. isotope event uc20b1, at 131 m, is characterised by a slight increase in δ13c values followed by a decrease of 0.2‰. uc20b2, at 118 m, is reflected by a change in δ13c values of 0.2‰. uc20b3 occurs immediately prior to the onset of marl deposition and is characterised by two short-term fluctuations in δ13c values of 0.2‰ each. the upper six events all post-date the marly interval. uc20c1, near the boundary between subzones uc20b and uc20c, is characterised by a drop in δ13c values and marks the termination of the early maastrichtian δ13c high. similarly uc20d1 is characterised by a marked fall in δ13c values immediately above the boundary between subzones uc20c and uc20d (fig. 3). isotope event uc20c2, at 72 m, is characterised by a change in δ13c values of 0.2‰. uc20c3, at 68 m, is reflected by a decrease in δ13c values of 0.3‰ and uc20c4 is characterised by a 0.3‰ increase in δ13c values. correlation to adjacent areas carbon isotope data also have been aquired from the maastrichtian chalk of the 250 m long karlslunde-1 core (eastern sjælland), from mid-maastrichtian chalk and marl from the rørdal quarry (fig. 3), and from upper maa strichtian chalk from the m-10x well in the danish north sea (figs 1, 3). 14 fig. 2. chalk and cyclical chalk–marl succession spanning the macrofossil defined lower–upper maastrichtian boundary exposed in the rørdal quarry. the exposed section consists of a lower homogeneous chalk unit, a middle cyclic chalk–marl unit and an upper unit of homogeneous chalk. the light beds are marl and the darker beds chalk. this ‘inversion’ of colour is due to surface weathering. the height of the section is 20 m. sedimentological and log data indicate that cyclically deposited chalk and marl are present in uc20b in both stevns-1 and karlslunde-1. the δ13c curve for karlslunde-1, although less densely sampled, displays the same overall pattern as the stevns-1 curve with an overall decrease in the δ13c values during the maastrichtian, starting with a stable early maastrichtian high followed by an interval with more fluctuating δ13c values (fig. 3). isotope events uc20c1-4 and uc20d1-2 are also easily identified in the karlslunde-1 succession whereas it has only been possible to identify the lowermost uc20b1 event in the interval below the chalk–marl cycles (fig. 3). the δ13c values in karlslunde-1 are approximately 0.1‰ higher than those in stevns-1. in the rørdal section it has been possible to identify isotope event uc20b3 immediately below an interval of interbedded chalk and marl of mid-maastrichtian age. at rørdal, uc20b3 is characterised by short-term changes of 0.2‰ in the 2–6 m interval (fig. 3). the δ13c values are approximately 0.2‰ more negative than those from the stevns-1 core. the upper maastrichtian chalk in the m-10x well was deposited in relatively deep water and differs from the timeequivalent onshore successions by the absence of flint nodules and layers (ineson et al. 2006). we have identified isotope events uc20c2, uc20c4 and uc20d1 in the cored interval (fig. 3). the uc20d1 event at 1990 m ends a long-term fall in δ13c values similar to those recognised in stevns-1 and karlslunde-1 (fig. 3). the absolute δ13c values are approximately 0.2‰ more negative than those recorded in stevns-1. isotope events uc20c2 and uc20c4 are both identified as short-term positive δ13c excursions of 0.1–0.2‰, similar to those seen in stevns-1 and karlslunde-1. variation in sedimentation rates the correlation of δ13c events across the danish area allows a refined stratigraphic subdivision of the chalk compared with the present nannofossil zonation, and thus leads to a more detailed model of sedimentation rates across the region. 15 m aa st ri ch tia n r ør da l, el ev at io n (m ) m -1 0x , d ep th ( m ) a b c d p. d an . lo w er up pe r n an no fo ss il zo ne st ag e g am m alo g g am m alo g g am m alo g d ep th ( m ) d ep th ( m ) u c 20 1 32 3 3 2 2 4 44 1 1 1 2 2 2 2 1 11 1 3 stevns-1 karlslunde-1 m-10x and rørdal 50 100 150 200 250 50 1960 1970 1980 1990 2000 2010 2020 20 10 0 –10 0 100 150 0.5cps 1.0 1.5 2.0 2.5 3.0 1 ? ? δ13c (‰) 0.5cps 1.0 1.5 2.0 2.5 3.0 δ13c (‰) 0.5cps 1.0 1.5 2.0 2.5 3.0 δ13c (‰) fig. 3. gamma-ray log and carbon isotope variation of the stevns-1, the karlslunde-1 and the m-10x (gamma-ray log not shown) and the rørdal section (gamma-log for stevns-1 obtained from core scanning (stemmerik et al. 2006)). the thick black vertical lines indicate recognised cyclic chalk–marl intervals. the full black horizontal lines represent correlation based on nannofossils (sheldon 2006). the stippled line shows the uncertain position of the nannofossil boundary. the thick grey lines represent running average of the data points. note the difference in vertical scale between the wells. 11 marked isotope changes out of the 29 changes occur in the depicted interval. p, paleocene; dan, danian; cps, counts per second. the present data set indicates that the accumulation rate in the karlslunde area was almost twice as high as at stevns-1 during subzone uc20c. the interval starting at uc20c1 and terminating at uc20d1 is 45 m thick in stevns-1 and 80 m thick in karlslunde-1. in both cores the sediment consists of flint-rich chalk. surprisingly, sedimentation rates for the upper maastrichtian in the north sea offshore area seem not to differ much from those recorded in the stevns-1 core. the interval from uc20c2 to uc20d1 is 25 m in m-10x and 35 m in stevns-1, and the thickness of the interval from uc20d1 to the maastrichtian–danian boundary is approximately 25 m in both areas. lateral variation of carbon isotopes the δ13c values in the cyclic chalk–marl interval in stevns-1 and karlslunde-1 show variations of up to 0.2‰, with a mean value of 2.3‰. in the rørdal quarry, the δ13c values in the time-equivalent interval also show variations of 0.2‰, but the mean value is 2.1‰. similarly, the average δ13c value of the chalk belonging to nannofossil subzone uc20c in m10x is 0.2‰ lower than that of the time-equivalent chalk in stevns-1 and karlslunde-1. this east–west, intra-basinal variation in δ13c values most likely reflects compositional differences of the chalk, being richer in macrofossils towards the east. conclusions the high-resolution carbon isotope curve constructed for the upper campanian to uppermost maastrichtian chalk in stevns-1 allows 29 isotope events to be identified. the re gional significance of eight of these events has been tested using isotope data from onshore and offshore cores and an outcrop. most isotope events are readily identified in the wells and sections though they are generally less densely sampled. it is evident that δ13c data can be used for correlation of chalk at nannofossil subzone level, and give a better understanding of lateral variations in sedimentation rates. acknowledgement the danish natural science research council supported the project financially. references ineson, j.r., buchardt, b., lassen, s., rasmussen, j.a., schiøler, p., schovsbo, n.h., sheldon, e. & surlyk, f. 2006: stratigraphy and palaeoceanography of upper maastrichtian chalks, southern danish central graben. geological survey of denmark and greenland bulletin 10, 9–12. jarvis, i., mabrouk, a., moody, r.t.j. & de cabrera, s. 2002: late cretaceous (campanian) carbon isotope events, sea-level change and correlation of the tethyan and boreal realms. palaeogeography, palaeoclimatology, palaeoecology 188, 215–248. sheldon, e. 2006: upper maastrichtian–danian nannofossils of the danish central graben and the danish basin: a combined biostratigraphic-palaeoecological approach, 354 pp. unpublished ph.d. thesis, university college london, uk. sheldon, e. in press: upper campanian calcareous nannofossil bio stratigraphy of the stevns-1 borehole, denmark. journal of nanno plankton research 30. stemmerik, l., surlyk, f., klitten, k., rasmussen, s.l. & schovsbo, n. 2006: shallow core drilling of the upper cretaceous chalk at stevns klint, denmark. geological survey of denmark and greenland bulletin 10, 13–16. surlyk, f. 1984: the maastrichtian stage in nw europe, and its brachiopod zonation. bulletin of the geological society of denmark 33, 217–223. 16 authors’ addresses n.h.s., s.l.r. & e.s., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: nsc@geus.dk l.s., department of geography and geology, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. geological survey of denmark and greenland bulletin 12, 73-77 73 1487.7 1623.5 freja mb lark fm 1400 1500 1600 1700 m frida-1 gr sonic ber in the inez-1 well (shown as unnamed sandstones in fig. 2); these sandstones may be contemporaneous or even contiguous with those of the freja member. however, confident correlation on the basis of log and seismic data is not possible at present. acknowledgements aage bach sørensen (geus) is thanked for help with seismic interpretations. yvonne desezar, johnny e. hansen and birthe amdrup are thanked for preparation of microfossil and palynology samples. the referees robert w.o’b. knox (british geological survey) and paul van veen (conocophilips norway) are thanked for their constructive criticism of the manuscript; the editorial contributions of adam a. garde, jon r. ineson and martin sønderholm are gratefully acknowledged. this work was made possible through grants from the danish energy authority, under the energy research project framework 2000. fig. 62. frida-1, reference well for the freja member. black bar shows cored section. references ahmadi, z.m., sawyers, m., kenyon-roberts, s., stanworth, c.w., kugler, k.a., kristensen, j. & fugelli, e.m.g. 2003: paleocene. in: evans, d. et al. 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(eds): petroleum geology of the continental shelf of north-west europe: proceedings of the 2nd conference, 3–39. london: heyden & son ltd. geological survey of denmark and greenland bulletin 17, 2009, 33-36 geomorphological indications of tectonic activity in the danish glacial landscape were pointed out already by milthers (1916, 1948). he described a conspicuous system of n–s-trending, narrow valleys in central jylland and interpreted them as fault-generated features (fracture valleys). the valleys occur in the area between ulstrup and hammel and in a smaller area near skjød (fig. 1). the most significant valley system is found near hvorslev, and it is here referred to as the hvorslev lineaments (fig. 1). an alternative interpretation of the genesis of the hvorslev lineaments was presented by hansen (1970). he argued that the n–s-trending valleys were formed by backward erosion from e–w-orientated erosional valleys in a former drainage system related to former higher groundwater table. however, milther’s interpretation that the valleys have a tectonic origin was later supported by larsen & kronborg (1994) and torp (2001). former interpretations of the lineaments were primarily based on morphological studies and arguments. recently the area with the hvorslev lineaments was mapped as part of the systematical geological mapping of denmark by the geolo gi cal survey of denmark and green land. the mapping of the surface li tho logy has added new information, and along with available seismic data it provides a frame work for a well-found ed interpretation of the elongated valleys, which is presented in this paper. we conclude that the fracture valleys are tectonic features, based on their morphology and because they are situated above a fault zone. geological setting the highest levels in the region are covered by a till unit, mainly a clayey till, but with patches of sandy till (fig. 2). on fracture valleys in central jylland – a neotectonic feature peter roll jakobsen and stig a. schack pedersen © geus, 2009. geological survey of denmark and greenland bulletin 17, 33–36. available at: www.geus.dk/publications/bull 33 bbjbjbjebjebjeeeebjeebjebjebjebjebbjbjebjebjebjebjebjejbjejbjebjebjebjebjeebjebjebjebjebjeebjebjeejebjbbbjjjjjj rrrrrrrrrrrrrrrrrrrrrrrrrrrrrrrrrrrrrrirrrrrrrrrrrrrrrrrrrrrrrrirrrrrrrrrrrirrrrrrrriririrr bngbngbbnggbgbngbbbbbbbngbngbbbngbngbngbgbbbbbbbbbbbngbgbbbbgbbbngbbbngbngbbbbngbnnngbn bngbngbngbbngbbngbngbngbngbngbgngbg roororororroroorrroroooorororoorrororoorororooroooroorobjerringbro hvohvohvohvohvohvohvovohvovohvohvohvohhhvohvohvohvohvohvovohvovohvohvohvohvoohvhvoovoovvvvvhvovvovhvvohvov rrrrrsrsrrsrsrsrrrsrsrssssrsrsrrrrrr vevevevevvevevevevevevvve uuluuuluuuuuululuuluuuuuululsu suuuulsuuulsulssuulsuulsuuuulsssstttttrutrutrtttrut uttrutrut uttrtrutrutrtrurrrutrtrtrutrutrtrutrurupppppppppppppppppppppppppppppppppppppp hhamhhahamhamhamhamhhhhamhamhamhamhamhamhhahamhamhamhamhhamhhamhamhhhammmhh mmmmemememmemememememmemmmm hinhinhhhhinhinhinininnnhinh nnhhh nh nhinhhhh nhhhhhhhhhhhhhhhhhhhhhhhinhhhhinnnernenenernenerenerererrreerneererererrrruuuupupupupupuupppuupupppuppp ttthothothothohothoththothothothothooootthothohoootthoothotththohotththoothtthohththhhh rrsrsrsørsrssøsørsørsørsørsørsørsrsøørsørsøøøørrrrr skjskjskjkjskjskjkskjkskjkjskjskjskskjskjskjskkskjskskskjskjskskjskjskjskjskjjskjkjjsk ødødødødødødødødødødødøddøddød udgudguddgudgududddudgudgududdguddgudgudguddggudggudgudguudggudddgudddddduddddududddenåenåenåenenåenenåennåenåeenåenåenåenenånånånåenånenånånenåenååenånenååennånånnåneenåenåååå hvorslev ulstrup hammel hinnerup thorsø skjød gudenå 5 km fig. 1. digital elevation model of the hvorslev region with place names mentioned in the text. the focus area is the northern part of the valley system between ulstrup and hammel near hvorslev. the difference in elevation from the valleys to the highest areas is around 90 m. the rectangle shows the location of fig. 3, and the inset map of denmark shows the location of the model. fig. 2. geological map of the hvorslev area (extract from the digital map, see jakobsen & hermansen (2007). contour interval 5 m. the location of the seismic line 73203 is shown with dots, which indicate the position of every 10th geophone. 60 m sandy till gravelly till clayey till gravel sand clay gravel sand postglacial freshwater deposits peat sand, clay extra-marginal deposits glaciofluvial deposits glacial deposits built-up areas2 km ulstrup bjerringbro hvorslev 1080 1100 11201090 1110 rosa_2008:rosa-2008 01/07/09 15:48 side 33 the steep slopes of the gudenå valley, glaciofluvial deposits are found below the till unit. smaller outcrops of glaciofluvial sand appear in the till unit representing erosional windows in the till surface. thus the uppermost lithostratigraphy in the region comprises a glaciofluvial unit overlain by a till. this succession is also recorded in boreholes in the region (gravesen 1991). in most wells a clayey till is present at the top below which the glaciofluvial sand has a thickness of up to 20 m. the sand rests mainly on oligocene mica clay and in a few places on oligocene quartz sand. the glaciofluvial unit is correlated with the teppestrup formation (larsen et al. 1977; pedersen & petersen 1997), which was deposited on an outwash plain in front of the ice advancing from the north-east. the clayey till is interpreted as deposited by the ice advance from the north-east that reached the main stationary line in jylland. the till is correlated with the mid danish till of houmark-nielsen (1987) and its correlative the fårup till of kronborg et al. (1990), which date to around 23 000–21 000 years before present in this part of denmark (houmark-nielsen & kjær 2003). morphology the parallel alignment of valleys in the ulstrup–hammel area suggests that they can be characterised as fracture valleys (fig. 1). the length of the valley system between ulstrup and hammel is about 17 km. the most prominent part of the system is located at hvorslev, where the valleys trend n–s with a spacing of 125 m to 250 m. the individual valleys are up to about 9 km long and 6 m deep. the width of the zone with fracture valleys is 800–1100 m wide (figs 1, 3). the valleys cut hills and erosional valleys without any change of direction and are therefore not controlled by the general landscape morphology. even one of the highest hills in the area is cut at the top by one of the valleys. the individual valleys show an undulating floor in the longitudinal direction, occasionally with small depressions containing bogs and lakes. most cross-sections of the individual valleys display a u-shaped morphology; in some places an asymmetrical shape is seen with a steep eastern slope and a gently dipping western slope. the trends of the westernmost valleys south of hvorslev are concave towards the west, curving around a depression which is filled with postglacial freshwater deposits. downhill towards this depression, and almost perpendicular to the hvorslev lineaments, a system of erosional valleys cuts into the surface with a gentle fall towards the west. the most pro minent erosional valley system is first order valleys with branching second order valleys. the major first order valley cuts through the hvorslev lineaments and a stream starts in its lower part where one of the n–s-trending valleys intersects the erosional valley. the second order valleys are not as deep as the first order valleys, and are cut by the hvorslev lineaments. some of them are hanging valleys on the sides of the n–s-trending valleys. the e–w-trending erosional valleys were generated shortly before the hvorslev lineaments. most of them are dry valleys that formed at a time when surface drainage was greater than now, most likely shortly after deglaciation. ero sion within the first order valleys was so strong that it sustained the effect of the formation of the hvorslev lineaments, which were apparently almost penecontemporaneous with the erosional valleys. seismic records the seismic section 73203 shows an e–w cross-section across the hvorslev lineaments south-east of bjerringbro (fig. 4). the features between geophone positions 1090 and 1112 are inter34 s n 500 m 50 m fig. 3. block diagram of the hvorslev lineaments. the inferred faults are drawn on the frontal south-facing side of the block diagram. the arrows mark some of the erosional valleys. 1080 1100 1120 1140 0 0.5 1.0 1.5 2.0 2.5 3.0 tw ow ay t ra ve l t im e (s ec ) geophone positions eeww ew base chalkbase chalk 2 km fig. 4. part of the seismic section 73203 in the hvorslev region (for location see fig. 2). the fault zone is situated between geophone positions 1090 and 1112, and the faults are marked with black lines. the height of the section corresponds to c. 3000 m. rosa_2008:rosa-2008 01/07/09 15:48 side 34 preted as faults. they occur over a zone c. 2.5 km wide situated directly below the hvorslev lineaments. the overall struc ture is interpreted as a negative flow er structure, where extension caused downfaulting of seg ments into a minor graben bounded by marginal faults. across the inferred fault zone downfaulting towards w predominates. the largest normal fault displacement is 0.05 second at 1.4 second two-way travel time measured on the re flec tor of the base chalk, which corresponds to a downthrow of c. 50 m at 1400 m depth. east of the fault zone the reflector is horizontal, and to the west it dips weakly towards w. indications of faulting at the surface most of the area around the hvorslev lineaments is covered by till. along the steepest slopes of the asymmetrical valleys, glaciofluvial sand is recorded below clayey till and on the gently dipping slopes clayey till drapes the surface all the way to the bottom of the valley (fig. 5). our interpretation of this distri bution of the lithological units is that the upper till unit has been displaced by a normal fault down to the floor of the valley, whereby the glaciofluvial sand is exposed in the footwall (fig. 5). the fault planes dip steeply w with a vertical displacement of c. 6 m. discussion the correlation between the hvorslev lineaments and the faults inferred from the seismic section indicates that the lineaments are surface traces of deep-seated faults. the hvorslev lineaments strike n–s, which most likely corresponds to the orientation of the faults. the himmerland graben is located to the north of the hvorslev lineaments (fig. 6; vejbæk 1990). it is a deep-seated structure outlined by two major n–s-trending basementattached faults, which is associated with the sorgenfrei– tornquist zone (fig. 6). dextral strike-slip faulting in the sorgenfrei–tornquist zone has resulted in e–w extension across the himmerland graben and led to normal faulting. the major tectonic events which caused syn-rift subsidence within the himmerland graben occurred during the triassic and in the late cretaceous. the hvorslev lineaments are situated south of the eastern fault of the himmerland graben and probably form a southern continuation of this fault. the indication of downthrow to wards w within the hvorslev lineaments corroborates this assumption. this implies that the fault activity in the him merland graben continues to the south into the central part of jylland. the fault-related valleys formed after deposition of the young est till in the area, because the till is cut by faults (fig. 5). the youngest glaciation of the area was during the main weichselian ice advance, during which the mid danish till was deposited. because the fracture valleys and the erosional valleys were formed at the same time, the faulting probably occurred shortly after the recession following the main weich selian ice advance. the deglaciation of the region occurred around 21 000 years ago, which gives a maximum age for the formation of the fracture-valley system in central jylland. two different causes can be suggested for the generation of the faults. either the fault activity was a response to the glaciostatic rebound after the last deglaciation, or the faulting is related to the neotectonic extension responsible for the general graben subsidence in the danish basin. this question cannot be answered unambiguously on the basis of the available data, but the trend of the lineaments and graben faults are oblique to the general trend of the postglacial marine limit in the region (mertz 1924) and hence are probably not linked to postglacial isostatic movements. the fact that the asymmetrical valleys show a downthrow towards w indicates that the valleys were formed in connection with tectonic activities in the him merland graben. moreover, the long length of the fault-related valleys also supports a tectonic origin. 35 fig 5. photograph of one of the asymmetric valleys (viewed southwards). in the westfacing, steep slope, glaciofluvial sand is present below a clayey till. the height of the steep slope is 6 m. on the gently dipping slope clayey till drapes the surface down to the bottom of the valley. the inserted block diagram illustrates the interpretation of the valley as generated by faulting. till till glaciofluvial sand c. 10 m w till glaciofluvial sand fau lt e rosa_2008:rosa-2008 01/07/09 15:48 side 35 in addition to the fracture-valley system between ulstrup and hammel, another system of similar valleys occurs near skjød, as pointed out by milthers (1916, 1948). fracture valleys are also found farther to the east near hinnerup (fig. 1; czakó 1994). morphologically they are similar to the hvorslev lineaments, although they are more curved. nevertheless, we suggest that they are fault-controlled. conclusions we suggest that the system of long, straight and narrow valleys in central jylland is related to a fault zone recognised on a seismic section across the valley system. hence the valleys are regarded as surface traces of a deep-seated fault zone. the faulting resulted in downfaulting of the surface-forming till along normal faults with downthrow mainly to w. the faulting occurred at a time when surface drainage was greater than at present, shortly after the last deglaciation that occurred around 21 000 years ago. the deformation is consequently a neotectonic activity. the fracture valleys indicate that the him merland graben continues farther to the south than previously outlined, and that crustal deformation related to this tectonic feature took place during the late quaternary. fault-related valleys as described in this paper are geomorphological elements that characterise some parts of the danish landscape. it is important to consider their potential implications for geological and hydrological models developed for the quaternary deposits. references czakó, t. 1994: the photogeological map and the map of surface water flow net of grundfør. explanations, preliminary results. dgu datadokumentation 3, 18 pp. gravesen, p. 1991: geological map of denmark, 1:50 000. map sheet bjerringbro. geological basic data map. danmarks geologiske under søgelse kortserie 32. hansen, k. 1971: de miltherske spaltedale i jylland. dansk geologisk forening, årsskrift for 1970, 47–53. houmark-nielsen, m. 1987: pleistocene stratigraphy and glacial history of the central part of denmark. bulletin of the geological society of denmark 36, 1–189. houmark-nielsen, m. & kjær, k. 2003: southwest scandinavia, 40–15 kyr bp: palaeogeography and environmental change. journal of quaternary science 18, 769–786. jakobsen, p.r. & hermansen, b. 2007: danmarks digitale jordartskort 1:25 000, version 3.0. danmarks og grønlands geologiske under søgelse rapport 2007/84, 27 pp. kronborg, c., bender, h., bjerre, r., friborg, r., jacobsen, h.o., kristiansen, l., rasmussen, p., sørensen, p.r. & larsen, g. 1990: glacial stratigraphy of east and central jutland. boreas 19, 273–287. larsen, g., jørgensen, f.h. & priisholm, s. 1977: the stratigraphy, structure and origin of glacial deposits in the randers area, eastern jutland. danmarks geologiske undersøgelse ii. række 111, 36 pp. larsen, g. & kronborg, c. 1994: geologisk set, det mellemste jylland. en beskrivelse af områder af national geologisk interesse. 272 pp. odense: geografforlaget. mertz, e.l. 1924: oversigt over de senog postglaciale niveau forandringer i danmark. danmarks geologiske undersøgelse ii. række 41, 49 pp. milthers, v. 1916: spaltedale i jylland. danmarks geologiske under søgelse iv. række 1(3), 16 pp. milthers, v. 1948: det danske istidslandskabs terrænformer og deres opstaaen. danmarks geologiske undersøgelse iii. række 28, 234 pp. pedersen, s.a.s. & petersen, k.s. 1997: djurslands geologi, 96 pp. copen hagen: geological survey of denmark and greenland. sigmond, e.m.o. 2002: geological map, land and sea areas of northern europe, scale 1:4 million. trondheim: geological survey of norway. torp, s. 2001: de miltherske spaltedale – landskabsdannelse og laserscanning i midtjylland. geologisk nyt 3(1), 28–29. vejbæk, o.v. 1990: the horn graben, and its relationship to the oslo graben and the danish basin. tectonophysics 178, 29–49. 36 authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: prj@geus.dk 56°30´n 10°e himmerland graben danish basin sorgenfrei–tornquist zone 25 km lineamentsfaults fig. 6. map of central jylland showing the tectonic outline of the region. fault-related valleys (red box) after milthers (1916) are located in the continuation of the eastern boundary fault of the himmerland graben. faults after vejbæk (1990) and sigmond (2002). rosa_2008:rosa-2008 01/07/09 15:48 side 36 geological survey of denmark and greenland bulletin 15, 2008, 57-60 57 the need for accurate predictions of future environmental change under conditions of global warming has led to a great interest in the most pronounced climate change known from the holocene: an abrupt cooling event around 8200 years before present (present = a.d. 1950), also known as the ‘8.2 ka cooling event’ (ka = kilo-annum = 1000 years). this event has been recorded as a negative δ18o excursion in the central greenland ice cores (lasting 160 years with the lowest temperature at 8150 b.p.; johnsen et al. 1992; dansgaard 1993; alley et al. 1997; thomas et al. 2007) and in a variety of other palaeoclimatic archives including lake sediments, ocean cores, speleothems, tree rings, and glacier oscillations from most of the northern hemisphere (e.g. alley & ágústsdóttir 2005; rohling & pälike 2005). in greenland the maximum cooling was estimated to be 6 ± 2°c (alley et al. 1997) while in southern fennoscandia and the baltic countries pollenbased quantitative temperature reconstructions indicate a maximum annual mean temperature decrease of around 1.5°c (e.g. seppä et al. 2007). today there is a general consensus that the primary cause of the cooling event was the final collapse of the laurentide ice sheet near hudson bay and the associated sudden drainage of the proglacial lake agassiz into the north atlantic ocean around 8400 b.p. (fig. 1; barber et al. 1999; kleiven et al. 2008). this freshwater outflow, estimated to amount to c. 164,000 km3 of water, reduced the strength of the north atlantic thermohaline circulation and thereby the heat transported to the north atlantic region, resulting in an atmospheric cooling (barber et al. 1999; clark et al. 2001; teller et al. 2002). the climatic consequences of this meltwater flood are assumed to be a good geological analogue for future climate-change scenarios, as a freshening of the north atlantic is projected by almost all global-warming models (e.g. wood et al. 2003; ipcc 2007) and is also currently being registered in the region (curry et al. 2003). in an ongoing project, the influence of the 8.2 ka cooling event on a danish terrestrial and lake ecosystem is being investigated using a variety of biological and geochemical proxy data from a sediment core extracted from højby sø, north-west sjælland (fig. 2). here we present data on changes in lake hydrology and terrestrial vegetation in response to climate change, inferred from macrofossil data and pollen analysis, respectively. materials and methods højby sø is located in odsherred, north-west sjælland, approximately 2.5 km from the sea. today the lake has a surface area of c. 40 ha with a mean water depth of 1.8 m. the lake has no natural inlets or outlets. in 2005 a 13.6 m long sediment © geus, 2008. geological survey of denmark and greenland bulletin 15, 57–60. available at: www.geus.dk/publications/bull environmental response to the cold climate event 8200 years ago as recorded at højby sø, denmark peter rasmussen, mikkel ulfeldt hede, nanna noe-nygaard, annemarie l. clarke and rolf d. vinebrooke hudson bay lake agassiz north atlantic ocean denmark fig. 1. map showing lake agassiz in north america and the route of the meltwater outburst into hudson bay and the north atlantic ocean when the lake drained at about 8400 b.p. (modified from kleiven et al. 2008). 58 core consisting mainly of calcareous gyttja was retrieved from the lake. from the entire core 28 samples were dated by accelerator mass spectrometry (ams) 14c using terrestrial plant material. the whole sediment sequence covers the time period c. 12,000–2000 years b.p.; here we focus on time slices of relevance to the 8.2 ka cooling event. the core content of plant and animal macrofossils is shown as accumulation rates. pollen data were calculated as percentages, concentrations (grains/cm3) and accumulation rates (grains/cm2 per year); only the accumulation rate data are presented here. at least 500 pollen grains from terrestrial plants were counted per sample. results and discussion lake hydrology and climate change sedimentary macrofossil data can be used as proxy evidence for changes in catchment and lake hydrology (e.g. hannon & gaillard 1997). at højby sø the abundance of macrophyte remains – ceratophyllum (hornwort), nymphaea (white water-lily), nuphar (yellow water-lily), najas marina (hollyleaved naiad), chara (stonewort) – and daphnia resting eggs (ephippia) exhibits an abrupt increase around 8400 b.p. (fig. 3). this pronounced change testifies to a sudden precipitation-induced lake level rise, as a higher water table would, on the one hand, result in an extension of shallow areas suitable for macrophyte growth, and on the other hand enhance conditions for the pelagic-living daphnia. this inferred change to moist conditions and increased lake level are supported by simultaneous and marked increases in sediment accumulation rates of minerogenic matter and the alga pediastrum (not illustrated). these data from højby sø add to the growing evidence that the brief 8.2 ka cooling event observed and defined in the greenland ice cores (8247–8086 b.p.; thomas et al. 2007) took place during a period of longer-term climatic perturbation which started some hundred years earlier (e.g. rohling & pälike 2005; lal et al. 2007). interestingly, the start of the moist climatic period inferred from højby sø (c. 8400 b.p.) is contemporary (within dating uncertainties) with the weakening of the thermohaline circulation (kleiven et al. 2008) and a global co2 decline of c. 25 ppmv as inferred from stomata analysis from lille gribsø, north-east sjælland (wagner et al. 2002). germany sweden 50 km 56°nhøjby søjylland sjælland fyn 56°n 10°e 10°e 7400 7600 7800 8000 8200 8400 8600 8800 9000 9200 9400 20 ce rat op hy llu m sp ., l ea f fr ag men ts nym ph ae a s p., se ed fr ag men ts nup ha r s p., se ed fr ag men ts nym ph ae ac ea e, tri ch os ch ler eid s naja s m ari na , s ee d f ra gm en ts ch ara sp ., o os po re s 20 dap hn ia, ep hip pia age (y ea rs b.p .) rad ioc ar bo n d ate d l ev els 1200 1250 1300 dep th be low w ate r s ur fac e ( cm ) 1350 macrofossil remains/cm2 per year fig. 2. map of denmark showing the location of højby sø in north-west sjælland. fig. 3. macrofossil accumulation diagram (remains/m2 per year) from højby sø covering the time period c. 9400–7400 years b.p. only selected taxa are shown. the age-depth model will be published in a forthcoming paper. 59 vegetation and climate change figure 4 illustrates the pollen accumulation rates for selected tree taxa and groups of taxa at højby sø in the time interval 8400–7400 years b.p. in this study period the overall trends in the pollen data are similar whether calculated as accumulation rates, percentages or concentrations. in the atlantic chronozone (9000–5900 years b.p.), the danish landscape was characterised by broad-leaved, closed-canopy woodlands, also called the ‘stable primeval forest’ (iversen 1973). within the dating uncertainty of our chronology the pollen strati graphy at højby sø provides clear evidence for vegetational disturbances coeval with the 8.2 ka cooling event. between c. 8250 and 7900 years b.p. there is a pronounced decline and subsequent recovery in the pollen accumulation rates for tilia (lime), quercus (oak), corylus (hazel) and alnus (alder). the beginning of the decline in each of the mentioned taxa and the subsequent recovery are as follows: tilia c. 8250/8100 b.p., quercus c. 8200/8000 b.p., corylus c. 8250/7900 b.p. and alnus c. 8100/8000 b.p. by contrast, ulmus (elm) accumulation rates are more variable and do not show a similarly clear decrease, although the taxon has a minimum frequency about 8100 b.p. two taxa, betula (birch) and pinus (pine), exhibit a clear maximum in the time interval c. 8100–8000 b.p. our data suggest that the primary response to the 8.2 ka cooling event was a decrease in the total pollen accumulation rates of thermophilous (‘warm-loving’), deciduous tree taxa in the time period c. 8200–8000 b.p. (fig. 4). this tree pollen recession is probably not a reflection of reduced forest cover as the abundance of open ground herbs – e.g. artemisia (mugwort), rumex acetosella (sheep’s sorrel) and poaceae (grasses) – does not exhibit a contemporary increase (not illustrated). as discussed by several authors, the decrease in pollen abundance of a number of thermophilous broad-leaved tree taxa during the 8.2 ka cooling event need not be synonymous with a change in population size. instead it might, solely or partly, represent reduced pollen production due to unfavourable climatic conditions (e.g. snowball et al. 2002; seppä et al. 2007). the various taxa referred to above do not respond simultaneously to changing environmental conditions, which might be due to differences in their physiological tolerance towards changes in, for example, temperature and hydrology. the decrease in pollen accumulation rates for tilia and quercus, which flower in july and may/june respectively, strongly suggests that the forest ecosystem in our study area was stressed by low temperatures during the summer season. many european palaeoclimate records and model simulations indicate that the temperature drop during the 8.2 ka cooling event was primarily a winter and early spring phenomenon (alley & ágústsdóttir 2005; wiersma & renssen 2006). thus, our findings at højby sø constitute one of the rare examples of the 8.2 ka cooling event also being a summer phenomenon. the decline in the pollen accumulation rates of the early flowering taxa corylus, alnus and ulmus (start flowering february–april) was most likely caused by long winters with late spring frosts that would have damaged flowers and catkins, leading to a reduction in pollen productivity. in creases in pollen accumulation rates for betula and pinus during the 8.2 ka cooling event are presumably due to the fact that these two taxa are the most frost-resistant tree taxa in northern europe. however, the fact that the accumulation rates for the latter taxa actually increase compared to levels the rm op hil ous tr ee s que rcu s til ia co ryl us aln us ulm us be tul a pin us 7400 7600 7800 8000 8200 8400 a ge ( ye ar s b. p. ) 20,000 60,000 1000 3000 5000 15,000 10,000 20,000 10,000 20,000 2000 6000 5000 10,000 2000 4000 pollen grains/cm2 per year fig. 4. pollen accumulation rates (grains/cm2 per year) of sum of thermophilous trees and dominant tree taxa at højby sø during the time interval c. 8400–7400 b.p. note different horizontal scales. the time period of the 8.2 ka cooling event according to the greenland ice core chronology is indicated by blue (8247–8086 b.p.; thomas et al. 2007). immediately before the cooling event, is cause for speculation. the elevated accumulation rates during the 8.2 ka cooling event might reflect an increase in actual population size and not just an increase in pollen productivity. if this is correct, the cause of the population expansion could be a hydrological change towards drier conditions. this hypothesis finds some support in the clear inverse relationship between the decrease in alnus and the increase in betula and pinus, suggesting a causal link between these taxa. alnus is a tree usually associated with damp or waterlogged soils and therefore sensitive to changes in the water table; the alnus decrease might therefore, solely or partly, be related to a water-level lowering in this period, and with an exposure of the former littoral zone around the lake betula and pinus may have expanded into this habitat. conclusions and future work due to good chronological control (not presented here) combined with high sampling resolution of the sediment sequence from højby sø, it has been possible to identify the 8.2 ka cooling event in a danish palaeo-record. the pollen data reveal that the forest ecosystem was affected by low temperatures during both the summer and winter – early spring; the result was reduced pollen production from thermo phil ous, deciduous trees. possible changes in population size due to climatic-induced hydrological changes are also suggested. furthermore, our investigation indicates that the short 8.2 ka cooling event took place during a period of longer-term climatic deterioration which started around 8400 b.p., coeval with the catastrophic drainage of lake agassiz. using diatom and algal pigment analyses, ongoing work in the højby sø project aims to explore if and how the aquatic ecosystem responded to the climate change 8200 b.p. acknowledgements geocenter copenhagen and the carlsberg foundation are gratefully acknowledged for financial support. references alley, r.b. & ágústsdóttir, a.m. 2005: the 8k event: cause and consequences of a major holocene abrupt climate change. quaternary science reviews 24, 1123–1149. alley, r.b., mayewski, p.a., sowers, t., stuiver, m., taylor, k.c. & clark, p.u. 1997: holocene climate instability: a prominent widespread event 8200 years ago. geology 25, 483–486. barber, d.c. et al. 1999: forcing of the cold event 8,200 years ago by catastrophic drainage of laurentide lakes. nature 400, 344–348. clark, p.u., marshall, s.j., clarke, g.k.c., hostetler, s.w., licciardi, j.m. & teller, j.t. 2001: freshwater forcing of abrupt climate change during the last glaciation. science 293, 283–287. curry, r., dickson, b. & yashayaev, i. 2003: a change in the freshwater balance of the atlantic ocean over the past four decades. nature 426, 826–829. dansgaard, w. 1993: evidence for general instability of past climate from a 250-kyr ice core record. nature 364, 218–220. hannon, g.e. & gaillard, m.-j. 1997: the plant-macrofossil record of past lake-level changes. journal of paleolimnology 18, 15–28. ipcc (intergovernmental panel on climate change) 2007: climate change 2007. (www.ipcc.ch). iversen, j. 1973: the development of denmark’s nature since the last glacial. danmarks geologiske undersøgelse v. række 7-c, 126 pp. johnsen, s.j., clausen, h.b., dansgaard, w., fuhrer, k., gundestrup, n., hammer, c.u., iversen, p., jouzel, j., stauffer, b. & steffensen, j.p. 1992: irregular glacial interstadials recorded in a new greenland ice core. nature 359, 311–313. kleiven, h.f., kissel, c., laj, c., ninnemann, u.s., richter, t.o. & cortijo, e. 2008: reduced north atlantic deep water coeval with the glacial lake agassiz freshwater outburst. science 319, 60–64. lal, d., large, w.g. & walker, s.g. 2007: climatic forcing before, during, and after the 8.2 kyr b.p. global cooling event. journal of earth system science 116, 171–177. rohling, e.j. & pälike, h. 2005: centennial-scale climate cooling with a sudden cold event around 8,200 years ago. nature 434, 975–979. seppä, h. et al. 2007: spatial structure of the 8200 cal yr bp event in northern europe. climate of the past 3, 225–236. snowball, i., zillén, l. & gaillard, m.-j. 2002: rapid early-holocene environmental changes in northern sweden based on studies of two varved lake-sediment sequences. the holocene 12, 7–16. teller, j.t., leverington, d.w. & mann, j.d. 2002: freshwater outbursts to the oceans from glacial lake agassiz and their role in climate change during the last deglaciation. quaternary science reviews 21, 879–887. thomas, e.r., wolff, e.w., mulvaney, r., steffensen, j.p., johnsen, s.j., arrowsmith, c., white, j.w.c., vaughn, b. & popp, t. 2007: the 8.2 ka event from greenland ice cores. quaternary science reviews 26, 70–81. wagner, f., aaby, b. & visscher, h. 2002: rapid atmospheric co2 changes associated with the 8,200-years-b.p. cooling event. proceedings of the national academy of sciences of the united states of america 99, 12,011–12,014. wiersma, a.p. & renssen, h. 2006: model-data comparison for the 8.2 ka bp event: confirmation of a forcing mechanism by catastrophic drainage of laurentide lakes. quaternary science reviews 25, 63–88. wood, r.a., vellinga, m. & thorpe, r. 2003: global warming and thermohaline circulation stability. philosophical transactions of the royal society a361, 1961–1974. authors’ addresses p.r., geological survey of denmark and greenland. øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: per@geus.dk m.u.h. & n.n.-n., department of geography and geology, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. a.l.c., apem manchester lab, riverview, a17 embankment business park, heaton mersey, stockport, sk4 3gn, uk. r.v., department of biological science freshwater biodiversity laboratory, university of alberta, edmonton, alberta t6g 2e9, canada. 60 madsen et al. 2024: geus bulletin 57. 8364. https://doi.org/10.34194/geusb.v57.8364 1 of 8 method article | short introducing inpox: a method for informed point extraction from geological 2d surfaces exemplified on the danish national hydrostratigraphic model rasmus bødker madsen1* , frederik alexander falk2 , ingelise møller1 , anne-sophie høyer1 1near-surface land and marine geology, geological survey of denmark and greenland (geus), aarhus, denmark; 2department of geoscience, aarhus university, aarhus, denmark abstract this study presents a probabilistic method for extracting informed points from geological surfaces, named inpox. the method generates a probability map from the existing surface by calculating the laplacian at each location and combining it with a user-defined transfer function. a set of points from the surface is then extracted with a density proportional to the probability map. the method allows a de-coupling of the most informative points in the surface from points carrying less or even biased information. inpox can be applied on any geological surface where the user needs to retrieve the structurally relevant parts and remove the information created by the initial interpolation. here, we test inpox on synthetic data, with and without supressing interpolation artifacts. in both cases, the informed points extracted with inpox outperforms a uniform probability map in recreating the original features. we show that the method requires a minimum of points to be extracted for inpox to be more informative than a uniform point retrieval. finally, to showcase the strength of the method in both retrieving the relevant geological features and suppressing the existing interpolation artifacts, we apply inpox to a real case surface from the danish national hydrostratigraphic model. *correspondence: rbm@geus.dk received: 11 dec 2023 revised: 25 feb 2024 accepted: 12 mar 2024 published: 27 may 2024 keywords: inpox, laplacian, decoupling information, interpolation artifacts, 3d layer models, artifact suppression abbreviations fohm: danish national hydrostratigraphic model (fælles offentlig hydrostratigrafisk model) inpox: informed point extraction lm: laplacian map pm: probability map rm: random map mad: mean absolute deviation geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: julian koch (geus, denmark) reviewers by: jacob skauvold (norwegian computing center, norway) volume: 57 funding: see page 7 competing interests: see page 8 additional files: none introduction in 3d geological modelling, models are either constructed as voxel models (van der meulen et al. 2013; madsen et al. 2021a; kawo et al. 2023) or layer models (lelliott et al. 2006; kaufmann & martin 2008; robins et al. 2008; wycisk et al. 2009; royse 2010; madsen et al. 2022). in the case of layer models, which are often used in subsequent hydrological modelling (seifert et al. 2012; enemark et al. 2023), geological units are defined by layer boundaries or surfaces, determined based on elevation on a 2d spatial grid. a common tabular abstract geographical coverage n/a temporal coverage n/a subject(s) covered computational geoscience, informatics and remote sensing method type the presented method is computational and is newly developed method name inpox instruments and equipment used • computer • programming language related publications n/a potential application(s) for this method the method can be applied to retrieve the most informed points from a given geological surface while downplaying the significance of interpolation artifacts. this allows a re-interpolation or simulation of the surface without including the information on the initial interpolation. https://doi.org/10.34194/geusb.v57.8364 https://orcid.org/0000-0001-8538-7491 https://orcid.org/0009-0003-1023-5161 https://orcid.org/0000-0002-1154-3700 mailto:rbm@geus.dk https://creativecommons.org/licenses/by/4.0/deed.ast madsen et al. 2024: geus bulletin 57. 8364. https://doi.org/10.34194/geusb.v57.8364 2 of 8 geusbulletin.org method for generating such layer models is through manual interpretation where a series of interpretation points are placed in space based on the available geological data or information and interpolation is used to estimate the surfaces between the interpreted points. thus, the chosen interpolation method and the way the geologist interacts with the interpolation method influence the characteristics of the final model and may also lead to interpolation artifacts on the surface (groshong 2006; yilmaz 2007; setianto & triandini 2015). modellers often rely on commercial software with built-in interpolation routines, which may lead to a lack of awareness regarding specific interpolation routines and parameterisation (wellmann & caumon 2018). this is true for both explicit and implicit geological modelling as some form of interpolation is a prerequisite in both modelling techniques. ultimately, the final product of a 3d geological modelling exercise will be the model itself. specifically for a layer model, the output is the layers, while the initial interpretation points and information about the interpolation are considered by-products and are in many cases lost. this makes it difficult, if not impossible, to redo the interpolation to test alternative parameterisations or to perform geostatistical simulation based on point information. it is therefore necessary to establish practical methods that can reverse engineer the process of interpolation to obtain a set of points from which the interpolation could be redone. the simplest way to sample from any surface is to draw a subset of the grid points using a uniform distribution. to solve issues of underrepresentation of some areas by the uniform point extraction, sampling methods have been developed that ensure a more even spatial coverage (brus et al. 2006; walvoort et al. 2010). within the scientific field of computer vision, the problem of surface recreation has been studied extensively with the purpose of rendering primary structural features and suppressing noise (see e.g. agrawal et al. 2006; boissonnat & oudot 2006; harker & o’leary 2015). however, to our knowledge no method or algorithm is currently available that considers the spatial properties of the surface when sampling points from a (geological) surface, although hansen (2021) and madsen et al. (2021b) showed that information content differs between points. thus, the selection of points matters in terms of recreating the essential features of the surface. in this short article, we address this problem and present a method, informed point extraction (inpox), for extracting points of high geological importance from a geological surface while suppressing interpolation artifacts. the method is based on quantifying the curvature by calculating the laplacian for the entire surface to assess its spatial properties and subsequently drawing points based on these values. the method is tested in a synthetic case to test the validity of inpox in a controlled setting as well as demonstrated in an applied case with a cut-out from a pre-existing 2d surface from the danish national hydrostratigraphic model (fohm; miljøstyrelsen 2023). resources required the following resources are required for this method: 1. a computer 2. a programming language 3. data in the form of (a) geological surface(s) methodological protocols inpox requires three steps: (1) construction of a laplacian map (lm), (2) a probability map (pm) based on a transfer function between probability and laplacians and (3) a random point retrieval based on the pm. these steps are explained in detail next. laplacian to inform the point extraction in inpox, the user must quantify the spatial characteristics of the surface. here, the laplacian is utilised to quantify the rate of change of the gradient in the surface, that is its curvature. the laplacian (∆) defines the divergence (∇) of the gradient (∇z) of the elevation (z) of the given surface as: l z x y z x y z x y x z x y y , . , , ,2 2 2 2( ) ( ) ( ) ( ) = ∆ = ∇ ∇ = ∂ ∂ + ∂ ∂ (1) where x and y are the two spatial coordinates in the grid. to compute the laplacian map, the operator is simply applied to all z in the geological surface producing a laplacian map, with higher values relating to areas with more variation in the surface. transfer functions to select areas with geologically relevant features, and to suppress the effect of unwanted simulation artifacts in the surface, a transfer function f(l) must be constructed that provides a probability drawing for  the current location (x,y) as a function of the laplacian. a basis probability level p0 is first set in inpox. p0 ensures that all areas of the surface have some probability of being drawn and that a desired fraction of the surface is drawn. for instance, if 5% of the surface should be drawn, p0= 0.05 is selected. in the case that, f(l) = p0, the result will be a random https://doi.org/10.34194/geusb.v57.8364 http://www.geusbulletin.org/ madsen et al. 2024: geus bulletin 57. 8364. https://doi.org/10.34194/geusb.v57.8364 3 of 8 geusbulletin.org selection from a uniform distribution. to manipulate the basis probability, a series of probability functions (f1 (l), f2 (l), f3 (l),…) can be introduced that either increases or decreases the basis probability depending on the laplacian. madsen et al. (2021b) and enemark et al. (2023) argue that major changes in the elevation of a geological surface in manual interpretation models are related to geologically relevant features. this is argued since variations in elevation mainly arise due to active choices made by the interpreter, for example, when mapping rivers, buried valleys and faults. we introduce f1(l), which increases the probability of drawing a point from the surface as a function of the laplacian l: f l cp l al g 2 tanh ( ) 11 0 0( )( ) = − + (2) where cp0 is the maximum probability added to p0 and thereby sets importance of the informed point retrieval in contrast to the base level. g is the maximum derivative of f1 (l) with respect to l. l0 is a reference laplacian and al0 is the value of l where the steepest rise of f1 (l) occurs. a and g thereby define how quickly and how abruptly f(l) becomes (1+c)p0 rather than p0 when l increases. a graphical representation of f1 (l) can be seen in fig. 1 (cyan shading, left hand side). mathematical interpolation of geological point information results in spatial continuous surfaces. hence, larger discontinuities can stem from either an active choice made by the modeller (e.g. introducing faults) or from the interpolation routines (e.g. using a limited neighbourhood of points locally instead of all points to make the algorithms more computationally feasible). in a system where there are no faults mapped or the laplacian is significantly different than the faults, a second function f2(l) can be designed to reduce the base probability and minimise the effect of interpolation artifacts: f l c d p l bl el dp ( ) ( ( ) ) 1 2 0 0 2 0 0( ) = + − + − (3) where dp0 is the maximum probability subtracted from p0 , and elo functions as a dampener going from (1+c) p0 to (1−d)p0. bl0 is the laplacian defining the switch between domain f1 (l) and f2 (l). a graphical representation of f2 (l) can be seen in the purple part of fig. 1. in inpox, the following transfer function is chosen as a combination of f1 (l) (eq. 2) and f2 (l) (eq. 3) with blo as the separation criteria between two: f l f l p l bl f l p l bl 1 0 0 2 0 0 ( ) ( ) ( )= + < + ≥     (4) (l >= bl0): f2(l) p0 bl0 (l < bl0): f1(l) pr ob ab ilit y laplacian (l) al0 el0 dp0 cp0 (1+c)p0 (1–d)p0 al0 bl0 (b+e)l0 g (b–a)l0 f(l) = f1(l) + f2(l) + p0 l0 fig. 1 proposed transfer function f (l) in inpox. the red line shows the probability (on the y-axis) as a function of ∆z (on the x-axis). the parameter b divides the values where f1 (l) and f2 (l) are used, shown by cyan and purple shading, respectively. other parameters are explained in the main text. https://doi.org/10.34194/geusb.v57.8364 http://www.geusbulletin.org/ madsen et al. 2024: geus bulletin 57. 8364. https://doi.org/10.34194/geusb.v57.8364 4 of 8 geusbulletin.org random point retrieval once a pm is constructed, the task of drawing the points is trivial. firstly, the pm is scaled such that the average probability of the entire grid becomes p0. a random map (rm) of similar size to the pm is constructed with values from an uncorrelated uniform distribution between 0 and 1. at each location where the rm value is lower than the pm value, a point is extracted from the surface grid. validation using synthetic data a synthetic reference (surface) model is created (fig. 2a) with different features that need to be emphasised in (a) reference model 0 25 50 75 100 y –40 –30 –20 –10 0 elevation (m a.s.l.) (b) laplacian map 0 25 50 75 100 0 5 10 15 20 laplacian 0 5 10 15 20 25 30 laplacian 0 0.05 0.1 0.15 0.2 0.25 pr ob ab ilit y (c) transfer functions t1 t2 t3 transfer functions (d) t1 probability map 0 25 50 75 100 y 0 0.05 0.1 0.15 0.2 0.25 probability (h) t2 probability map 0 25 50 75 100 y (l) t3 probability map x 0 25 50 75 100 y 0 25 50 75 100 (e) t1 draw: 489 points 0 25 50 75 100 (i) t2 draw: 502 points x 0 25 50 75 100 (m) t3 draw: 498 points (f) t1 interp. (mad = 1.1214) 0 25 50 75 100 (j) t2 interp. (mad = 1.2368) 0 25 50 75 100 (n) t3 interp. (mad = 1.3885) x 0 25 50 75 100 (g) t1 repeat (mad = 1.0592) 0 25 50 75 100 (k) t2 repeat (mad = 1.1046) 0 25 50 75 100 (o) t3 repeat (mad = 1.2278) x 0 25 50 75 100 t1 t2 t3 0 10025 50 75 0 10025 50 75 0 10025 50 75 0 10025 50 75 0 10025 50 75 0 10025 50 75 0 10025 50 75 0 10025 50 75 0 10025 50 75 0 10025 50 75 0 10025 50 75 0 10025 50 75 0 10025 50 75 0 10025 50 75 fig. 2 results from the synthetic case study. a: the reference model. b: the corresponding laplacians. c: the transfer functions (t1, t2 and t3) used in the experiment. results are arranged by transfer functions. t1: emphasises all large laplacians. t2: emphasises all large laplacians while downplaying the square. t3: uses a uniform distribution. each row contains: a probability map (d, h, l); a single draw of point extraction (e, i, m); a linear interpolation of the drawn points (f, j, n), 100 repeats of drawing 500 points and interpolating (g, k, o). the mean absolute deviation (mad) is shown for each interpolation. on average 500 points are drawn from the three probability maps, but the number of points drawn in a single realisation varies as shown in panels e, i and m. https://doi.org/10.34194/geusb.v57.8364 http://www.geusbulletin.org/ madsen et al. 2024: geus bulletin 57. 8364. https://doi.org/10.34194/geusb.v57.8364 5 of 8 geusbulletin.org the point extraction. at x = 50 and y = 75, a square is placed that is highly discontinuous with the surrounding surface, while the other features are less discontinuous. the lm is calculated from the reference model and is presented in fig. 2b. figure 2c shows two different transfer functions (t1 and t2) to translate the lm into pms as well as a transfer function where f(l) = p0 (red dotted line) that enables a comparison to random point retrieval (t3, orange line). for t1 (blue line), we treat all changes in the surface as relevant information and emphasise the full spectrum of laplacians found in fig. 2b. for t2 (red line), we treat the highly discontinuous square as an interpolation artifact, and only values < 10 are emphasised while values > 10 are suppressed. the parameterisation of each transfer function is presented in table 1. three pms are obtained by applying the three transfer functions. the results are shown in fig. 2d, h, l, where it is clearly visible that changes in elevation are reproduced using the inpox strategy and that different points are subsequently drawn based on these probabilities (fig. 2e, i, m). in the third column (fig. 2f, j, n), the points are linearly interpolated to create a new surface. the biggest difference between t1 and t2 and the uniform draw (t3) is the lack of dark red and black areas in t3 and thus a lack of structural information and a smoother output surface. between t1 and t2, the square is more clearly defined in t1 as significantly more points were drawn to help define its shape. the mean absolute deviation (mad) is calculated in all cases. the mad confirms the visual validation as t1 and t2 interpolations are closer to the reference model than a uniform draw. thus, the extracted points carry more of the structurally important information for recreating the reference model. the mad for t2 is higher than t1 because the suppression of the square extends the difference to the reference model at that location. to ensure that these results did not arise because of a fluke draw, we repeat the experiment of drawing a set of points from the pm and interpolating between these 100 times to form a series of new surfaces. the average surface for each transfer function (displayed in fig. 2g, k, o) confirms that t1 most accurately recreates the reference model on average, while t2 recreates the reference model while downplaying the significance of the square. to assess the effect of the number of points being drawn in inpox, we repeat the experiment, generating 100 surfaces from 100 set of points at various basis draw probabilities (here denoted as the draw percentage) of the total grid size. figure 3 shows the results for draw percentages in the interval between 1% and 13% of the surface being extracted, including the experiment shown in fig. 2g, k, o that corresponds to a draw percentage approximating 5%. in general, as the draw percentage increases, the informed point retrievals from t1 and t2 decrease the mad significantly more than t3. at approximately 3%, there are similar levels of mad, and for very sparse data extraction (< 3%), the pattern is reversed, such that the uniform distribution outperforms t1 and t2. because the informed extraction tends to cluster the drawn points, some areas may end up with a very low density of points when only a few points are extracted, which makes it difficult to recreate the surface in these areas. in contrast, the uniform extraction always secures an evenly spatial distribution of drawn points and is more robust in cases where data extraction is sparse. naturally, these numbers vary depending on the choice of reference model and transfer function and should not be used directly as decisive indicators of where inpox is a relevant point-extraction strategy. instead, in a broader sense, these results indicate that inpox can be a relevant strategy when a certain number of points need to be extracted, but some minimum of point draw percentage is needed for the method to be effective. table 1 transfer-function parameters and basis probability (draw percentage) for each of the transfer functions used in the study. t1–t3 are used in the synthetic case study, while t4 is used in the real-world case. transfer function p0 [−] l0 [m -1] a[−] b[−] c[−] d[−] e[−] g[m] t1 0.05 1 1.5 25 4 1 1 2 t2 0.05 1 1.5 6.5 4 1 2 2 t3 0.05 t4 0.02/0.20 1 1.5 3 5 1 0.8 2 0 2 4 6 8 10 12 ave. draw percent age (%) 0.5 1 1.5 2 2.5 3 3.5 4 m ad (1 00 d ra w s) mad vs. draw percentage t1 t2 t3 fig. 2 experiment fig. 3 mad for 100 point extractions and subsequent interpolations calculated as a function of draw percentage for the three transfer functions t1, t2 and t3. the experiment run in fig. 2 is shown as a vertical black dotted line. https://doi.org/10.34194/geusb.v57.8364 http://www.geusbulletin.org/ madsen et al. 2024: geus bulletin 57. 8364. https://doi.org/10.34194/geusb.v57.8364 6 of 8 geusbulletin.org demonstration using a real-world application we apply inpox on a surface from the danish national hydrostratigraphic model (fohm; fælles offentlig hydrostratigrafisk model). this model was created by correlating and interpolating across multiple local models. this has led to many interpolation artifacts that reside amongst valuable geological features. this is exemplified in fig. 4a through an excerpt of a quaternary surface from the fohm model covering 40.4 km2. a river-like structure passes through an area of the surface that is clearly affected by interpolation artifacts. these artifacts can be seen as round shapes (marked with arrows) due to interpolation distances set in the interpolation algorithm. we apply inpox to extract points from the surface with the input parameters shown in table 1 as t4. from the lm (fig. 4b), we assess that most of the geologically relevant features have laplacians in the interval of 2 and 4, while the artifacts typically have laplacians larger than ten but can go down to values around 5. thus, we set b = 3 and c = 5p0 to emphasise the features and e = 0.8 to ensure that all laplacians above 5 are downplayed in significance. in fig. 4c, d, we show the result from a single draw with a draw percentage of 2% and 20%, respectively. they likely represent two extremes in terms of point extraction considering that the total number of interpretation points divided by the surface extent is a 9.9% coverage. in the case of the 2% extraction, the greatest concentration of points occurs along the river-like structure and near the highest elevation peaks in the terrain. conversely, point density is minimal in the low-lying areas, particularly to the south and west. for the 20% draw, the river-like structure is almost completely drawn out, while the artifacts are still difficult to detect visually. this demonstrates that inpox does exactly as intended in a real-world setting. (a) fohm-model layer 6.14 6.145 6.15 6.155 yu tm 106 –60 –40 –20 0 20 40 60 elevation (m a.s.l.) (b) laplacian map 5.35 5.4 5.45 5.5 xutm 105 6.14 6.145 6.15 6.155 yu tm 106 0 10 15 20 laplacian (c) probability map (2%) 0 0.2 0.4 0.6 0.8 1 probability (d) probability map (20%) 5.35 5.4 5.45 5.5 xutm 105 (e) 2% points drawn 5.35 5.4 5.45 5.5 xutm 105 (f) 20% points drawn int. artifacts geol. feature 5 fig. 4 results from the real-world case study. a: an excerpt from a surface in the danish national hydrostratigraphic model. b the corresponding laplacian map. c: 2% draw from the surface using inpox. d: 20% draw from the surface using inpox. e: random point extraction using p0 = 0.02. f: random point extraction using p0 = 0.20. geol.: geological. int.: interpolation. https://doi.org/10.34194/geusb.v57.8364 http://www.geusbulletin.org/ madsen et al. 2024: geus bulletin 57. 8364. https://doi.org/10.34194/geusb.v57.8364 7 of 8 geusbulletin.org discussion and outlook in this study we have suggested a method, inpox, based on calculating the laplacians at all locations and defining a suitable transfer function to emphasise and suppress certain structures in the surface. thus, inpox, to some extent, makes it possible to reverse engineer the process of geological interpolation by extracting the relevant geological features and making it possible to re-interpolate the information. we have successfully validated and demonstrated the applicability of the method in both a theoretical and practical case where geological features are identified and are able to guide the point extraction while interpolation artifacts are removed. obviously, inpox cannot recreate the exact location of points placed by the geologist during interpretation and thus some of the interpolation information cannot be removed using inpox. by setting the draw percentage, the user can find a suitable trade-off between (1) removing all interpolation information while removing some geological information or (2) keeping all geological information while keeping some interpolation information. if some geologically relevant points are available from for example borehole information, it is trivial to include these in the draw by setting the draw probability in the pm to 1. madsen et al. (2021b) suggest using the gradient as a measure to calculate the points that are most relevant for describing the surface. as was the case for the laplacians, the user would look for areas with a steep gradient. although the gradient will likely provide a reasonable indication of geologically relevant areas of the model, it is difficult to keep information on sharp transitions in the surface. our tests, although not shown here, confirm this, as the mad is always worse when using the gradient approach rather than the laplacian approach. since the laplacian is the divergence of the gradient, it is low where the gradient is steep, and high adjacent to these areas. this is more logical because it is more informative to have points where a surface structure begins and ends rather than having the point information on the decline and then recreating the start and end-points with interpolation. the design of the transfer function matters for the resulting point extraction. in the current setup, inpox uses a baseline probability that is manipulated with two separate transfer functions, one with the ability to enhance probability and one being able to decrease probability. as demonstrated, one can easily manipulate the parameters to fit the surface features of interest and problems at hand. furthermore, since all steps in the algorithm have negligible execution times, even on large surfaces, one can fine-tune the parameters of the transfer function in an iterative process until a satisfactory point-extraction strategy is found. for larger surfaces, it might be necessary to construct region-specific transfer functions, however one could also apply one global conservative transfer function that only suppresses extreme laplacian values. it is important to stress that inpox is not limited either by the specific parameterisation or the applied transfer function. one can freely choose another more specific or intricate transfer function if it provides a mapping between the laplacian and draw probability. for instance, and not shown here, we adapted a piecewise linear function mimicking the t2 transfer-function to a degree that very similar mad results could be obtained. to conclude, the practitioner should complete a problem-specific assessment of the (spatial) distribution of laplacians and then make an informed decision on the transfer-function that makes most sense geologically. finally, our results on the effect of the draw percentage indicate that choosing inpox over any other possible point extraction method, as exemplified with the uniform distribution, follows the well-known ‘no-free-lunch’ theorem (mosegaard 2012; shalev-shwartz & ben-david 2014). this states that the effectiveness of inpox will be problem dependent. thus, in some scenarios it could be better to implement another point extraction strategy. if the surface is rapidly fluctuating between neighbouring points, the entropy is high and the information content between points is little to non-existent. in this case, no point extraction method would be preferable. in general, the basic requirement for inpox is that spatial correlations exist in the surface and furthermore that the surface curvature relates to geological information or artifacts. we see two main scenarios where inpox can be a most valuable tool: (1) when dealing with national models, where quality control is often limited to assessing the behaviour of the surface at local scale through cross-sections or possibly 2d maps of the surface and (2) in areas where models are created using different strategies or operate at different scales depending on the resources available for model creation. under these circumstances, inpox can be a valuable tool to homogenise what the points represent between different models and even within a single model. acknowledgements the authors wish to thank geocenter denmark and the project decode-3d for providing funding for writing the manuscript. the method development was carried out in a consultancy project (nret 2022–2024) conducted for the danish environmental protection agency (miljøstyrelsen). additional information funding statement the method was developed during consultancy work for the danish epa, while the hours for turning the results into a scientific contribution and writing the manuscript was provided by geocenter denmark through the project decode3d. https://doi.org/10.34194/geusb.v57.8364 http://www.geusbulletin.org/ madsen et al. 2024: geus bulletin 57. 8364. https://doi.org/10.34194/geusb.v57.8364 8 of 8 geusbulletin.org authors’ contribution rbm: conceptualisation, software, investigation, writing – original draft preparation faf: conceptualisation, methodology, software, investigation, writing – reviewing and editing im: conceptualisation, supervision, investigation, writing – reviewing and editing ash: project administration, supervision, writing – reviewing and editing competing interests the authors declare no competing interests. additional files none provided references agrawal, a., raskar, r. & chellappa, r. 2006: what is the range of surface reconstructions from a gradient field? in: leonardis, a., bischof, h. & pinz, a. 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https://doi.org/10.1016/j.cageo.2007.09.001 https://doi.org/10.1016/j.cageo.2007.09.001 https://doi.org/10.1002/esp.1473 introducing inpox: a method for informed point extraction from geological 2d surfaces exemplified on the danish national hydrostratigraphic model introduction resources required methodological protocols laplacian transfer functions random point retrieval validation using synthetic data demonstration using a real-world application discussion and outlook acknowledgements additional information references figures fig. 1 proposed transfer function f (l) in inpox. the red line shows the probability (on the y-axis) as a function of δz (on the x-axis). the parameter b divides the values where f1 (l) and f2 (l) are used, shown by cyan and purple shading, respectively. other parameters are explained in the main text. fig. 2 results from the synthetic case study. a: the reference model. b: the corresponding laplacians. c: the transfer functions (t1, t2 and t3) used in the experiment. results are arranged by transfer functions. t1: emphasises all large laplacians. t2: emphasises all large laplacians while downplaying the square. t3: uses a uniform distribution. each row contains: a probability map (d, h, l); a single draw of point extraction.... fig. 3 mad for 100 point extractions and subsequent interpolations calculated as a function of draw percentage for the three transfer functions t1, t2 and t3. the experiment run in fig. 2 is shown as a vertical black dotted line. fig. 4 results from the real-world case study. a: an excerpt from a surface in the danish national hydrostratigraphic model. b the corresponding laplacian map. c: 2% draw from the surface using inpox. d: 20% draw from the surface using inpox. e: random point extraction using p0 = 0.02. f: random point extraction using p0 = 0.20. geol.: geological. int.: interpolation. tables table 1 transfer-function parameters and basis probability (draw percentage) for each of the transfer functions used in the study. t1–t3 are used in the synthetic case study, while t4 is used in the real-world case. geological survey of denmark and greenland bulletin 19, 2009, pp. 171 + 3 plates geological survey of denmark and greenland bulletin 19 � 2009 lithostratigraphy of the cretaceous–paleocene nuussuaq group, nuussuaq basin, west greenland gregers dam, gunver krarup pedersen, martin sønderholm, helle h. midtgaard, lotte melchior larsen, henrik nøhr-hansen and asger ken pedersen geological survey of denmark and greenland ministry of climate and energy geological survey of denmark and greenland bulletin 19 keywords lithostratigraphy, nuussuaq group, cretaceous, paleocene, west greenland, nuussuaq basin. cover illustration sedimentary succession of the nuussuaq group at paatuut on the south coast of nuussuaq, one of the classical localities for sedimentological and palaeontological studies. the photograph shows deep incision of the paleocene quikavsak formation into the upper cretaceous atane formation. the conspicuous red coloration is due to self-combustion of carbonaceous sediments. the upper part of the succession comprises volcanic rocks of the west greenland basalt group. the height of the mountains is c. 2000 m. photo: martin sønderholm. frontispiece: facing page view down into the narrow paatuutkløften gorge on the southern coast of nuussuaq, where coarse-grained, pale sandstones of the paleocene quikavsak formation fill a major incised valley cut into interbedded mudstones and sandstones of the cretaceous atane formation. sea-fog often invades the coastal valleys but typically dissipates during the day. photo: finn dalhoff. chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, geological institute, university of aarhus; minik rosing, geological museum, university of copenhagen; finn surlyk, department of geography and geology, university of copenhagen scientific editor of this volume: jon r. ineson editorial secretaries: jane holst and esben w. glendal referees: j. christopher harrison (canada), robert knox (uk) and t. christopher r. pulvertaft (denmark) illustrations: jette halskov digital photographic work: benny m. schark layout and graphic production: annabeth andersen printers: rosendahls . schultz grafisk a/s, albertslund, denmark manuscript submitted: 22 april 2008 final version approved: 1 september 2009 printed: 28 december 2009 issn 1604-8156 isbn 978-87-7871-260-8 citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 19, 171 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk or at www.geus.dk/publications/bull © de nationale geologiske undersøgelser for danmark og grønland (geus), 2009 for the full text of the geus copyright clause, please refer to www.geus.dk/publications/bull abstract . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . preface . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . introduction. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . previous work. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . the period before 1938 – the pioneers on the fossil floras . . . . . . . . . . . . . . . . . . . . . the nûgssuaq expeditions 1938–1968 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . the early hydrocarbon and coal-related studies 1968–1982 . . . . . . . . . . . . . . . . . . . . recent investigations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . geological setting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . nuussuaq group. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . kome formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . slibestensfjeldet formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . upernivik næs formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . atane formation. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . skansen member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . ravn kløft member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . kingittoq member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . qilakitsoq member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . itivnera bed . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . itilli formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . anariartorfik member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . umiivik member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . kussinerujuk member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . aaffarsuaq member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . kangilia formation. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . annertuneq conglomerate member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . oyster–ammonite conglomerate bed . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . quikavsak formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . tupaasat member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . nuuk qiterleq member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . paatuutkløften member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . agatdal formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . eqalulik formation. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . abraham member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . atanikerluk formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . naujât member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . akunneq member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . pingu member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . umiussat member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . assoq member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . references . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . appendix: place names and localities . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5 7 9 11 11 15 17 17 21 27 36 46 49 55 64 67 71 77 80 82 89 92 96 99 104 110 112 114 118 121 123 127 133 138 139 146 149 152 154 155 159 160 169 contents 4 5 dam, g., pedersen, g.k., sønderholm, m., midtgaard, h.h., larsen, l.m., nøhrhansen, h. & pedersen, a.k. 2009: lithostratigraphy of the cretaceous–paleocene nuussuaq group, nuussuaq basin, west greenland. geological survey of denmark and greenland bulletin 19, 171 pp. the nuussuaq basin is the only exposed cretaceous–paleocene sedimentary basin in west greenland and is one of a complex of linked rift basins stretching from the labrador sea to northern baffin bay. these basins developed along west greenland as a result of the opening of the labrador sea in late mesozoic to early cenozoic times. the nuussuaq basin is exposed in west greenland between 69°n and 72°n on disko, nuussuaq, upernivik ø, qeqertarsuaq, itsaku and svartenhuk halvø and has also been recorded in a number of shallow and deep wells in the region. the sediments are assigned to the more than 6 km thick nuussuaq group (new) which underlies the palaeogene plateau basalts of the west greenland basalt group. the sediment thickness is best estimated from seismic data; in the western part of the area, seismic and magnetic data suggest that the succession is at least 6 km and possibly as much as 10 km thick. the exposed albian–paleocene part of the succession testifies to two main episodes of regional rifting and basin development: an early cretaceous and a late cretaceous – early paleocene episode prior to the start of sea-floor spreading in mid-paleocene time. this exposed section includes fan delta, fluviodeltaic, shelfal and deep marine deposits. the nuussuaq group is divided into ten formations, most of which have previously been only briefly described, with the exception of their macrofossil content. in ascending stratigraphic order, the formations are: the kome formation, the slibestensfjeldet formation (new), the upernivik næs formation, the atane formation (including four new members – the skansen, ravn kløft, kingittoq and qilakitsoq members – and one new bed, the itivnera bed), the itilli formation (new, including four new members: the anariartorfik, umiivik, kussinerujuk and aaffarsuaq members), the kangilia formation (including the revised annertuneq conglomerate member and the new oyster–ammonite conglomerate bed), the quikavsak formation (new, including three new members: the tupaasat, nuuk qiterleq and paatuutkløften members), the agatdal formation, the eqalulik formation (new, including the abraham member), and the atanikerluk formation (new, including five members: the naujât, akunneq (new), pingu (new), umiussat and assoq (new) members). abstract authors’ addresses g.d., dong energy, agern allé 24–26, dk-2970 hørsholm, denmark. e-mail: greda@dongenergy.dk g.k.p., department of geography and geology, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: gunver@geol.ku.dk m.s., geological survey of denmark and greenland. present address: dong energy, agern allé 24–26, dk-2970 hørsholm, denmark. e-mail: mason@dongenergy.dk h.h.m., dong energy, agern allé 24–26, dk-2970 hørsholm, denmark. e-mail: helmi@dongenergy.dk l.m.l., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: lml@geus.dk h.n.-h., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: hnh@geus.dk a.k.p., natural history museum of denmark, university of copenhagen, øster voldgade 5–7, dk-1350 copenhagen k, denmark. e-mail: akp@snm.ku.dk 6 7 the onshore cretaceous–paleocene sedimentary succession of the nuussuaq basin in west greenland has been studied since the mid-1850s, mainly because of its extremely well-preserved macroplant and invertebrate fossils and the presence of coal. the early work suggested major age differences between the various lithological units, but as knowledge increased as a result of the investigations carried out from the late 1930s to the late 1960s, the time gaps gradually diminished. during this period, several more or less informal stratigraphic schemes evolved and were even used differently by different authors, causing confusion concerning the actual definition, the vertical and lateral extent, and the age of the lithological units. this confusion was not helped by the fact that spelling conventions regarding greenlandic place names changed considerably over time. no formal lithostratigraphy in the nuussuaq basin was defined, although troelsen (1956) in his overview on stratigraphic units in greenland treated the units he described as formal. in the early 1990s, focus was again directed towards the region as an analogue for the offshore basins of west greenland. the danish state, and later the greenland government, provided substantial funding for studies that could counter the general assessment of the region as being only gas-prone. both extensive acquisition of seismic data and substantial onshore field studies were initiated in the nuussuaq basin during this period under the auspices of the geological survey of greenland (from 1995 the geological survey of denmark and greenland, geus). during the following years, the nuussuaq basin evolved from being an analogue for the offshore areas to being an exploration target in itself due to the finds of widespread oil seeps in the basin resulting in the drilling of the first onshore exploration well in greenland in1996. this culminated in 2007 when petroleum exploration offshore disko took a major leap forward with the granting of seven new exploration licenses. it is therefore evident that a formal description of the thick and varied succession onshore is strongly needed in order to create a common reference for geoscientists working in the region. the authors have contributed to various extents in the completion of the manuscript. gregers dam, gunver krarup pedersen and martin sønderholm have had dual roles as authors and compilers. they have provided original data on most of the formations, have written the introductory chapters and have supplied the majority of the figures. they have been responsible for the manuscript in all stages and have revised the manuscript in accordance with the comments from the referees. helle h. midtgaard has provided sedimentological logs and original observations on the kome, slibestensfjeldet and upernivik næs formations and on the ravn kløft member of the atane formation. henrik nøhr-hansen has examined numerous palynological slides and has provided data on the ages of most of the formations. lotte melchior larsen and asger ken pedersen have made it possible to correlate the siliciclastic sediments of the atanikerluk formation to the co-eval magmatic rocks, and have documented the areal extent of sedimentary and volcanic rocks on maps and vertical sections. preface ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ 200 km greenland nuussuaq basin baffin bay melville bay davis strait labrador sea 60°w 50°w 60°w 50°w70°w 74°n 70°n 66°n 66°n 70°n c d n s greenland possible oceanic crust or attenuated continental crust mesozoic basin palaeogene volcanic rocks shallow or exposed continental basement exploration wells exploration wells 1976–1977 faults■ fig. 1. simplified geological map of west greenland (for location, see inset) showing the nuussuaq basin in its regional setting, broadly outlined by the red box. based on escher & pulvertaft (1995), chalmers & pulvertaft (2001), oakey (2005) and gregersen et al. (2007). c, cape dyer; d, disko; n, nuussuaq; s, svartenhuk halvø. 8 the nuussuaq basin belongs to a complex of basins that were formed in the early cretaceous, extending from the labrador sea in the south to melville bay in the north (fig. 1; chalmers & pulvertaft 2001). due to local uplift during the neogene, it is the only one of these basins that extends into the onshore area in west greenland where upper cretaceous – paleocene sediments overlain by volcanic rocks can be studied in the disko – nuussuaq – svartenhuk halvø area. the nuus suaq basin has therefore been used for many years as an analogue for the basins offshore southern and central west greenland, the primary target for petroleum exploration. the former lithostratigraphy of the nuussuaq basin was established by researchers working with the classic flora and fauna of the nuussuaq basin, and definition of lithostratigraphical units was therefore to a large extent governed by fossil finds. this resulted in an incomplete lithostratigraphical scheme comprising some ill-defined units which were not true lithostratigraphic units. during the last two decades, the sedimentology, bio stratigraphy, sequence stratigraphy and organic geointroduction nuussuaq hareøen 70°n 55°w 51°w 51° w55°w uummannaq qaarsut vaigat malig aat disko ilulissatqeqertarsuaq aaffarsuaq ataata kuua saqqaqdalen saqqaq itil li skansen naassat ikorfat nuuk q iterl eq ataniker luk nuuk killeq assoq pingu qullissat qorlortorsuaq skarvefjeld sorte hak stordal fp93-3-1 uiffaq kussinerujuk kuugannguaq a lia na at su nn gu aq ilu gis so q asuk fig. 21 fig. 74 fig. 82 fig. 113 fig. 40 fig. 65 fig. 44 fig. 124 fig. 22 paatuut tupaasat eiq ak disko bugt 25 km upernivik øupernivik ø qeqertarsuaqqeqertarsuaq nuussuaq 72°n 71°n 70°n 69°n 51°w55°w inland ice vaigat disko 50 km grønne ejland aasiaat svartenhuk halvø ubekendt ejland upernivik ø qeqertarsuaq fig. 73 fig. 33 west greenland basalt group precambrian basement nuussuaq group town locality fig. 2. map of disko and nuussuaq showing the main localities and place names used in this paper. frames and figure numbers indicate coverage of detailed maps and a seismic section. the pre-quaternary geology is simplified from escher (1971). the detailed maps of upernivik næs (fig. 33) and svartenhuk halvø – qeqertarsuaq (fig. 73) are located on the regional map on the right. ak, akuliarusinnguaq; eiq, eqip inaarsuata qaqqaa. 9 fig. 3. topographical map by rink (1857) reproduced by nordenskiöld (1871), showing the location of kome, atane and other named features given lithostratigraphic significance by nordenskiöld. 10 the period before 1938 – the pioneers on the fossil floras the cretaceous–paleocene sediments of west greenland have been known as far back as the time when the norsemen in greenland visited this area and named a headland ‘eysunes’ from the norse word ‘eisa’ (meaning glowing embers); this refers to the spontaneous combustion of organic-rich mudstones and coal that has taken place after landslides at several localities along vaigat, the strait between disko and nuussuaq (fig. 2; rosenkrantz 1967; henderson 1969). following re-colonisation in 1721, the coal seams along vaigat again attracted much attention, and by the time the region had its first inspector in 1782 the disko bugt colonies had become self-sufficient in coal, as mentioned in paul egede’s ‘efterretninger om grönland’ (steenstrup 1874). the early geological and geographical investigations of this region were reported by giesecke (1806–13, in: steenstrup 1910), rink (1853, 1857; fig. 3), nordenskiöld (1871, 1872) and brown (1875). an account of the later investigations up to 1968 can be found in rosenkrantz (1970). for many years it was plant fossils found in the limnic part of the succession that attracted the attention of geologists from all over the world. brongniart (1831 p. 351) described the first species from kome on northern nuussuaq. in the following years, collections of plant fossils were made at several localities in the region, mainly from eastern disko and atanikerluk, which were described by heer (1868). these descriptions aroused so much interest that a british expedition led by e. whymper and r. brown was sent out to collect new material in 1867 (heer 1870). the expedition of a.e. nordenskiöld in 1870 provided a much larger collection, and for the first time upper cretaceous strata were recognised (heer 1874a, b). collections from an expedition in 1871 led by e.g.r. nauckhoff were described by heer (1880). important new collections were made by k.j.v. steenstrup during his expeditions in 1871–72 and 1878–80, and members of steenstrup’s expeditions discovered a number of new plant fossil localities on disko, nuussuaq, upernivik ø and svartenhuk halvø (fig. 4). steenstrup brought his collection back to copenhagen in 1880 and it was described by heer (1882, 1883a, b). heer (1883b) concluded from his studies, which now included more than 600 species (e.g. fig. 5), that the plant fossils could be divided into three cretaceous floras (the kome, atane and patoot floras) and one flora of tertiary age (the upper atanikerdluk flora). later work on plant fossils from this area includes that of seward (1924, 1926), miner (1932a, b; 1935), seward & conway (1935, 1939), koch (1963, 1964, 1972a, b), and boyd (1990, 1992, 1993, 1994, 1998a, b, c, 2000). the plant fossils from upernivik næs, described as being part of the atane flora by heer (1883b), were recognised as a separate flora (upernivik næs flora) by koch (1964) and referred to as the upernivik flora by boyd (1998a, b, c, 2000). previous work chemistry of the nuussuaq basin have been studied in detail by the survey and the university of copenhagen, and it is now possible to establish a modern, formal, lithostratigraphic framework for all the sedimentary units in the nuussuaq basin. as now defined, the individual cretaceous – early paleocene formations are, to a large extent, genetic units bounded by unconformities. the new framework has been established with the least possible alteration of the earlier defined units in order to avoid confusion and to promote the overall understanding of the basin. therefore, the naming of units does not in all cases conform to the rules set out by the north american commission on stratigraphic nomenclature (1983). some of the old informal units were named according to their content of fossils or lithology; a few of these units have been retained due to the large collections of fossils originating from them. the place names used herein are all in modern greenlandic orthography. however, previously defined lithostratigraphic units using older spelling have not been renamed. a complete list of place names used is given in both new and old orthography in the appendix; their location is shown on fig. 2. 11 fig. 4. the first geological map of the nuussuaq basin published by steenstrup (1883b). note that only the coastal areas were known as all travel was by boat. steenstrup made several long journeys in an umiaq, a traditional open rowing boat of seal-skin with a crew of women. 12 fig. 5. plant fossils from the atane formation at atanikerluk illustrated in flora fossilis arctica by heer (1883b). 13 in spite of the fact that this part of greenland had been visited by a large number of expeditions with geological objectives, knowledge of the marine strata was rather poor until the nûgssuaq expeditions from 1938 to 1968, probably because the marine fossils were overshadowed by the very well-preserved plant fossils. some marine fossils were collected in the latter part of the 19th century by g.f. pfaff, c.f.v. henriksen and greenlanders from niaqornat, and important collections were made by k.j.v. steenstrup during his expeditions in 1871–72 and 1878–80. additional sampling was carried out by d. white and c. schuchert in 1897 and by j.p.j. ravn and a. heim in 1909. steenstrup (1874) recognised the presence of marine strata within the atane formation of nordenskiöld (1871) along the south coast of nuussuaq (figs 4, 6). the collection of marine fossils made by steenstrup in 1871–72 in this area was examined by schlüter (1874 pp. 30–31), who concluded that the marine strata in west greenland must be of late cretaceous age. this conclusion was supported by de loriol (1883) who studied steenstrup’s 1879 collection of marine faunas from the north coast of nuussuaq and from paatuut and ataa on the south coast. the collection of scaphites from niaqornat, which was augmented considerably by greenlanders who accompanied steenstrup on his expeditions of 1878–80, was examined by v. madsen (1897) who referred the dominant species to the european senonian scaphites römeri d’orb. in 1897, a geological expedition under the auspices of the united states national museum in connection with the peary arctic expedition collected fossils from the cretaceous and tertiary localities on nuussuaq. stanton (in: white & schuchert 1898) was of the opinion that the collection of marine fossils from the north coast of nuussuaq included a number of characteristic late cretaceous types that could be equated to the senonian peak 1580 m peak 1722 m peak 1010 m tupaasat ivisaannguit ataata kuua ivissussat qaqqaat fig. 6. the south coast of nuussuaq around the ataata kuua river delta (the large valley in the centre of the profile) which is the type area of the nuussuaq group. upper panel: part of modern section measured by multimodel photogrammetry using oblique aerial photographs (from a.k. pedersen et al. 1993). lower panel: part of a long section painted by harald moltke, who accompanied k.j.v. steenstrup on his journey in 1898 (from steenstrup 1900). the 33 km long profile extends a little more to the west than the map in fig. 40. the profile shows many of the characteristic features of the nuussuaq basin: cretaceous sediments including the atane formation, up to 800 m (pale yellow); paleocene incised valleys, marine and lacustrine mudstones (grey), and hyaloclastic foreset-bedded breccias overlain by subaerial lava flows above c. 800 m (red, blue, purple and grey, hyaloclastite breccias in pale colours, lava flows in deep colours). outcrops of the quikavsak formation, including the type section, are shown in bright yellow below the summit point 1580 m (upper panel). 14 of europe, but he left the possibility open that some of the faunas from the south coast of nuussuaq could be tertiary in age. ravn (1918) arrived at the same conclusion as schlüter, de loriol and stanton concerning the age of the marine strata in west greenland, namely that they are senonian (late cretaceous). his study was based on old material stored in the mineralogical museum in copenhagen and collected by pfaff, henriksen and steenstrup together with material collected in 1909 by ravn, and to a lesser extent by heim, from disko and nuussuaq (heim 1910; ravn 1911). the work of ravn (1918) was at that time the most thorough study of the marine fossils and many species were illustrated for the first time. according to ravn (1918 p. 330) the occurrence of american upper cretaceous species in the greenland fauna suggested that a marine connection between west greenland and the central part of canada and the united states possibly existed in late cretaceous time, particularly as no affinities with the american coastal plain cretaceous fauna or the cretaceous of east greenland were evident. relying solely on ravn’s results, teichert (1939 p. 155) came to a similar conclusion, and in an accompanying map teichert showed the transgression of the late cretaceous sea to west greenland to be from the north-west. frebold (1934) described an early senonian fauna from east greenland and also made some remarks on the senonian fauna from west greenland. the pioneering stratigraphic papers by heer (1868, 1870, 1874a, b, 1880, 1882, 1883a, b), nordenskiöld (1871, 1872), and steenstrup (1883b) resulted in the establishment of three pre-volcanic lithostratigraphic units (the kome, atane and upper atanikerdluk formations) and four biostratigraphic units (the kome, atane, patoot and upper atanikerdluk floras). the first geological map of the region also derives from this period (fig. 4; steenstrup 1883b). the nûgssuaq expeditions 1938–1968 in 1938, the danish nûgssuaq expeditions ushered in a new epoch in the study of the cretaceous–tertiary sediments in west greenland. one of the many objectives of these expeditions was the study of the marine strata on nuussuaq and in other parts of the basalt region of west greenland (rosenkrantz 1970) and to obtain a more precise age for the limnic beds and their famous floras. in 1938 and 1939, the two nûgssuaq expeditions led by a. rosenkrantz and supported by the carlsberg foundation and the royal greenland trade department (den kongelige grønlandske handel) carried out studies in the nuussuaq basin. this work was continued after the second world war under the auspices of the newly established geological survey of greenland peak 2010 m peak 1900 m peak 1760 m peak 1580 m giesecke monument uppalluk 15 (grønlands geologiske undersøgelse, ggu). from 1948 to 1968, 16 expeditions to the area (14 led by a. rosenkrantz and two by s. floris and k. raunsgaard pedersen) had to a greater or lesser degree been involved in the study of the marine strata. a summary of the expeditions and their results was given by rosenkrantz (1970). more than 47 individuals participated in the work on the marine deposits during this 30-year period, some of whom are seen in fig. 7. these expeditions represented the first attempt to study systematically the marine cretaceous–tertiary sediments in west greenland. a new biostratigraphy was erected (rosenkrantz 1970 fig. 2) and it is thanks to the efforts of the members of the nûgssuaq expeditions that large collections of marine fossils were brought back to the geological museum in copenhagen. the fossils from these outstanding collections have since been described by many other workers: ammonites and belemnites: birkelund (1956, 1965) corals: floris (1967, 1972) coccoliths: perch-nielsen (1973), jürgensen & mikkelsen (1974) crustaceans: collins & wienberg rasmussen (1992) fish: bendix-almgreen (1969) foraminifera: h.j. hansen (1970) gastropods and bivalves: yen (1958), kollmann & peel (1983), petersen & vedelsby (2000) leaves and fruits: koch (1959, 1963, 1964) ostracods: szczechura (1971) palynomorphs: k.r. pedersen (1968) wood: mathiesen (1961) the gastropods were the subject of special study by rosenkrantz, but at the time of his death in 1974 only a fraction of the material had been published. a catalogue comprising the gastropod material left unpublished at rosenkrantz’s death was published by kollmann & peel (1983) and represents the culmination of many years of work by rosenkrantz and the technicians and artists under his direction. a catalogue of 115 bivalve taxa from the rosenkrantz collection has been published by petersen & vedelsby (2000). the gastropod family psedolividae has been revised by pacaud & schnetler (1999). however, large parts of the collection including echinoderms, serpulids, bryozoans, nautiloids, cirripeds, brachiopods and insects are still unpublished. aspects of the stratigraphy of the cretaceous–tertiary sediments in the disko – nuussuaq – svartenhuk halvø area were published in a number of papers between 1959 and 1976, in particular by koch (1959, 1963, 1964), koch & pedersen (1960), rosenkrantz & pulvertaft (1969), h.j. hansen (1970), rosenkrantz (1970) and henderson et al. (1976). some elements of the strati graphy presented in these papers are well established, especially the biostratigraphy of the marine cretaceous based on ammonites collected in situ (birkelund 1965) and the lithostratigraphy of the non-marine paleocene in southern nuussuaq (koch 1959). however, other stratigraphic interpretations were based on very general descriptions and a number of undocumented statements and correlations. a shortcoming of the work on the sediments during the nûgssuaq expeditions was the lack of a formal correlative, regional framework for the marine and non-marine cretaceous strata. fig. 7. members of the nûgssuaq expedition in 1949. back row (from left): johansi, abraham løvstrøm, sonja alfred hansen, andreas tobiassen, maggie graffpetersen. front row (from left): alfred rosenkrantz, kristian schou, bruno thomsen, christian poulsen, søren floris and eske koch. many of those in the photo accompanied rosenkrantz on several expeditions. rosenkrantz acknowledged the assistance of the hunters from niaqornat, and sonja hansen’s discovery of the most fossiliferous lithology when he established the andreas, sonja and abraham members of his agatdal formation. 16 17 a lithostratigraphy for the marine danian beds of northern and central nuussuaq was established by rosenkrantz (in: koch 1963) and rosenkrantz (1970). the paleocene kangilia and the agatdal formations were subdivided into members by rosenkrantz (1970), but these subdivisions are not entirely satisfactory, and rosenkrantz’s correlation of the marine paleocene across nuussuaq at member level was not documented. as regards the marine cretaceous, the situation was rather better thanks largely to the comprehensive work of birkelund (1965). she demonstrated that the marine cretaceous spanned the upper turonian – maastrichtian interval, and that the oldest marine strata are to be found on svartenhuk halvø. two map sheets in ggu’s 1:500 000 map series covering the central part of west greenland were compiled during this period (a. escher 1971; j.c. escher 1985). the early hydrocarbon and coal-related studies 1968–1982 in the late 1960s, the hydrocarbon potential of the basins offshore labrador and west greenland was recognised. in the early 1970s, petroleum exploration started with the acquisition of a large number of seismic surveys and culminated with the drilling of five exploration wells offshore west greenland in 1976 and 1977 (fig. 1; rolle 1985). in acknowledgement of the need for geological information, ggu and the petroleum industry studied the sedimentology, biostratigraphy and organic geochemistry of the onshore cretaceous–tertiary successions. however, these studies came to an abrupt end when all five wells drilled offshore were declared to be dry and the industry left west greenland. the most important results of these studies were published by henderson (1969, 1973, 1976), sharma (1973), elder (1975), schiener (1975, 1976), j.m. hansen (1976), henderson et al. (1976, 1981), schiener & floris (1977) and schiener & leythaeuser (1978). during this phase of the work, the first overall facies pattern was established and a palaeogeographic reconstruction for the mid-cretaceous was presented (schiener 1975, 1977), but neither a basin model nor systematic biostratigraphy or lithostratigraphy were published, despite the great efforts of both ggu and the petroleum industry. however, unpublished reports by ehman et al. (1976), croxton (1976, 1978a, b) and ehman (1977) and an unpublished thesis by j.m. hansen (1980b) presented new stratigraphic divisions and correlations based on palynology. the reports of ehman et al. and croxton dealt with the entire sedimentary succession of the region, but only in rather general terms. in contrast, j.m. hansen focussed on the marine paleocene and presented a more detailed stratigraphic analysis. a sedimentological paper on the lower cretaceous fluviodeltaic sediments at kuuk, on the north coast of nuussuaq was published by pulvertaft (1979). three 1:100 000 map sheets were compiled during this period (rosenkrantz et al. 1974, 1976; pulvertaft 1987). the first seismic lines in the nuussuaq basin were also acquired during this period (sharma 1973; elder 1975). according to the interpretations of these authors, the thickest succession of sediments is found along the north coast of nuussuaq, where it amounts to 4 km, of which 3 km are below sea level. along southern nuussuaq, the total thickness of sediments was estimated to be about 3 km, of which 2 km were interpreted to occur below sea level. following the abandonment of the coal mine at qullissat on the north-east coast of disko in 1972, a major study on the coals on nuussuaq was initiated by ggu in 1978 with support from the danish ministry of commerce. the aim of the project was to produce detailed geological and technical information on the coal-bearing strata on the south coast of nuussuaq, requisite information for a decision whether or not to invest in detailed exploration programmes (shekhar et al. 1982). three stratigraphic boreholes were drilled and a total of 828 m of core was taken. more than 12 km of outcrop section was measured and numerous coal samples were analysed. unfortunately, the results were disappointing and did not warrant further investigation; the results were never considered in more detail and were only documented in an internal project report (shekhar et al. 1982). although the study provided a lot of new information on the coals and the sedimentary succession, this information was never synthesised. recent investigations in the 1980s, petroleum geological research at ggu was focussed on north and east greenland and only limited research was carried out in the nuussuaq basin, mainly by geologists from the universities of copenhagen and aarhus (johannessen & nielsen 1982; g.k. pedersen 1989; boyd 1990, 1992, 1993, 1994, 1998a, b, c, 2000). it was not until the start of ggu’s disko bugt expeditions (1988–1992; fig. 8) that extensive studies of the cretaceous–tertiary sedimentary and volcanic succession of the nuussuaq basin were initiated together with studies on the adjacent precambrian terrain and studies relevant for mineral exploration in the region (kalsbeek & christiansen 1992, and references therein). during the first years of the disko bugt expeditions, most of the research on sediments dealt with the midcretaceous deltaic deposits of the atane formation and the synvolcanic lacustrine deposits. it was carried out mainly by the university of copenhagen and resulted in a number of case studies on sedimentology and palynostratigraphy (see below). the introduction of multimodel photogrammetry during this period (a.k. pedersen & dueholm 1992) permitted detailed mapping of individual rock units (e.g. a.k. pedersen et al. 1993, 2002). this provided a framework for later correlation of the synvolcanic sediments and interpretation of their palaeogeography (e.g. a.k. pedersen et al. 1996; g.k. pedersen et al. 1998). from the onset of field studies in the 1960s and the subsequent drilling of the five dry exploration wells in the mid-1970s, the lack of documented oil-prone source rocks in west greenland had been acknowledged as the main risk in relation to oil exploration in west greenland. in order to confront this problem, ggu initiated a reevaluation of the hydrocarbon potential. a systematic analysis of the nuussuaq basin was initiated in 1990 as part of the disko bugt expeditions including sedimentological, palynostratigraphical, source rock and diagenetic studies. at the same time, ggu reassessed some of the seismic data acquired during the early 1970s offshore west greenland, and it became evident that sedimentary basins which could contain oil and gas were much more extensive than had previously been supposed (chalmers 1989, 1993). this was further substantiated when seismic traverses across the labrador sea acquired in 1977 by bundesanstalt für geowissenschaften und rohstoffe (bgr) were reprocessed and reinterpreted; from the reprocessed lines it could be seen that the continent–ocean boundary lies much farther from the greenland coast than previously supposed, opening up the possibility that prospective sedimentary basins exist in the deeper parts of greenland waters (chalmers 1991; chalmers et al. 1993). this resulted in government-funded acquisition of more than 6000 km of seismic data by ggu in 1990–92 and a speculative survey of nearly 2000 km in 1992, all in offshore west greenland. in addition, more than 4000 km of seismic data in melville bay offshore north-west greenland were acquired by the industry in 1992. a licensing round was announced in west greenland in 1993, but no applications were submitted, mainly due to the lack of a documented oil-prone source rock in west greenland. the perception of the prospectivity of the area was significantly enhanced, however, by the discovery of bitumen in vugs in basalts near the base of the lava pile in 1992 and in a slim-core well (marraat-1) in 1993 (chri stiansen & pulvertaft 1994; dam & christiansen 1994). then, in 1994, a 13 km long reflection seismic line acquired on the southern shore of nuussuaq revealed that the base of the sedimentary basin is at least 5 km below sea level at this locality (christiansen et al. 1995). with these results, attention began to focus on the petroleum potential of the nuussuaq basin itself. after the original discoveries, oils and bitumen were found in surface outcrops over a wide area of western nuussuaq and also on the north side of disko and on the south-eastern corner of svartenhuk halvø (christiansen et al. 1998). inspired by the early finds, grønarctic energy inc., a small fig. 8. members of the disko bugt expeditions in 1992. studies on the sediments of the nuussuaq basin were carried out from a base in uummannaq. participants (left to right): ib olsen, søren saxtorph, gregers dam, eskild schack pedersen, helle h. midtgaard, christian j. bjerrum, lotte melchior larsen, asger ken pedersen, stig schack pedersen, martin sønderholm, christian schack pedersen, gunver krarup pedersen and flemming getreuer christiansen. 18 canadian company, held a concession in western nuus suaq from 1994 to 1998. during this period, the company drilled four slim-core wells (ganw#1, gane#1, gank#1 and gant#1) and one conventional exploration well (gro#3) to a depth of 2996 m (fig. 9). the wells were drilled in areas where geological information on the sediments beneath the volcanic rocks was completely lacking, and they provided a large volume of new data that also tied the outcrop areas together. detailed sedimentological, organic geochemical and palynostratigraphical analyses of the wells were carried out by the survey and were published in numerous reports of the ‘danmarks og grønlands geologiske undersøgelse rapport’ series in 1996–97. many of the results have later been incorporated into and summarised in other publications. the positive results from the years 1990–1994 encouraged the government of greenland and the danish state to provide further funding for studies to overcome the disappointing outcome of the 1992 licensing round. more field work was carried out on disko, nuussuaq and svartenhuk halvø in the period between 1994 and 2002. during the summer of 1995, the survey acquired more than 3700 km of seismic and gravity data, mainly in the fjords and sounds around disko and nuussuaq (chri stiansen et al. 1996a), and a 1200 m deep stratigraphic slim-core hole (umiivik-1) was drilled on svartenhuk halvø in 1995 (bate & christiansen 1996). additional information was obtained using seismic and magnetic data from the 1970s combined with an aeromagnetic survey flown over the area in 1997 (rasmussen et al. 2001). the most important results from these studies include: 1) interpretation of seismic and magnetic data, forward modelling of gravity profiles and a reappraisal of all available data on faults in the onshore areas (chalmers et al. 1999). 2) documentation of several oil types in surface oil seeps and wells and several possible source rock intervals and their possible correlation with cenomanian– turonian oils from the central western interior seaway in north america, or to upper jurassic-sourced oils from the jeanne d’arc basin offshore newfoundland and in the north sea (bojesen-koefoed et al. 1999, 2004, 2007; christiansen et al. 2002). 3) documentation of wet gas in the umiivik-1 well (fig. 73), suggesting the presence of a good, but overmature, turonian source rock for condensate or oil (dam et al. 1998b). 4) establishment of a new biostratigraphic scheme for the cretaceous and paleocene onshore and offshore deposits (nøhr-hansen 1996; nøhr-hansen et al. 2002; søn derholm et al. 2003). 5) documentation of promising reservoir intervals in the turbidite succession and a quantitative log-interpretation of the upper part of the gro#3 well (which was not tested prior to casing), suggesting high hydrocarbon saturations in sandstone units (kristensen & dam 1997; dam et al. 1998d; kierkegaard 1998). 6) completion of the geological mapping of eastern disko and south-east nuussuaq (a.k. pedersen et al. 2000, 2001, 2007a, b). extensive photogrammetrical work forms the basis of five geological profiles through the nuussuaq basin (1:20 000). these profiles document the sediments of the nuussuaq group and the overlying west greenland basalt group, including the relationships between the early volcanic rocks and the synvolcanic sediments (a.k. pedersen et al. 1993, 2002, 2003, 2005, 2006a, b). fig. 9. the gro#3 exploration well on western nuussuaq drilled by the canadian company grønarctic inc. in 1996. for location, see fig. 65; the drilled succession is shown in fig. 67. photo: kim zinckjørgensen. 19 the onshore studies carried out since 1988 have resulted in numerous publications, as summarised below: biostratigraphy and palaeontology: boyd (1990, 1992, 1993, 1994, 1998a, b, c, 2000), hjortkjær (1991), piasecki et al. (1992), koppelhus & pedersen (1993), nøhr-hansen (1993, 1996, 1997a, b, c), nøhr-hansen & dam (1997), dam et al. (1998b, c), kennedy et al. (1999), lanstorp (1999), nøhrhansen & sheldon (2000), nøhr-hansen et al. (2002), sønder holm et al. (2003), g.k. pedersen & bromley (2006). diagenesis: preuss (1996), stilling (1996), kierkegaard (1998). organic geochemistry: christiansen et al. (1996b, 1998, 1999, 2000), bojesenkoefoed et al. (1997, 1999, 2001, 2004, 2007), nytoft et al. (2000), g.k. pedersen et al. (2006). sedimentology: g.k. pedersen & jeppesen (1988), g.k. pedersen (1989), pulvertaft (1989a, b), midtgaard (1991), olsen (1991), olsen & pedersen (1991), g.k. pedersen & pulvertaft (1992), dueholm & olsen (1993), olsen (1993), dam & sønderholm (1994), a.k. pedersen et al. (1996), midtgaard (1996a, b), dam & sønder holm (1998), g.k. pedersen et al. (1998), dam et al. (1998a, 2000), jensen (2000), dam & nøhr-hansen (2001), dam (2002), nielsen (2003). structural geology: chalmers et al. (1999), j.g. larsen & pulvertaft (2000), chalmers & pulvertaft (2001), marcussen et al. (2002), bonow (2005), japsen et al. (2005, 2006, 2009), wilson et al. (2006), bonow et al. (2006a, b, 2007). well descriptions: dam & christiansen (1994), christiansen et al. (1994a, 1996c, 1997), bate & christiansen (1996), dam (1996a, b, c, 1997), dahl et al. (1997), kristensen & dam (1997), nøhr-hansen (1997a, b, c), kierkegaard (1998), ambirk (2000), madsen (2000). 20 as a result of the opening of the labrador sea in late mesozoic to early cenozoic times, a complex of linked rift basins stretching from the labrador sea to northern baffin bay developed along west greenland (fig. 1; chalmers & pulvertaft 2001). two main episodes of regional rifting and basin development during this time have been documented in the area: an episode of early cretaceous rifting, and a late cretaceous – early paleocene rift episode prior to the start of sea-floor spreading in mid-paleocene time (dam & sønderholm 1998; dam et al. 2000; chalmers & pulvertaft 2001; dam 2002; sørensen 2006). the most extensive outcrops of mesozoic–palaeogene rocks in the entire labrador sea – davis strait – baffin bay region are those of the nuussuaq basin in the disko – nuussuaq – svartenhuk halvø area in central west greenland. this basin may be a southern extension of the basin complex in the melville bay region (fig. 1; whittaker et al. 1997); the offshore area between 68° and 73°n is, however, covered by palaeogene basalts and little is therefore known about the deeper-lying successions in this region. a small outcrop is known from cape dyer, eastern baffin island (burden & langille 1990) and outcrops of cretaceous–palaeogene sediments are also seen farther north in arctic canada on bylot island (miall et al. 1980; miall 1986; harrison et al. 1999) and on ellesmere island (núñez-betelu 1994, núñez-betelu et al. 1994a, b; harrison et al. 1999). during the early paleocene (danian), the area offshore southern west greenland was subjected to major uplift and erosion (bonow et al. 2007). sedimentation resumed in the late danian contemporaneously with the major episode of paleocene volcanism in the disko–nuussuaq area and continued into the holocene with a major hiatus spanning the oligocene in the north and the mid-eocene to mid-miocene in the south (dalhoff et al. 2003). geological setting nuussuaq 72°n 71°n 72°n 51°w53°w56°w 71°n 70°n 69°n 55°w 51°w svartenhuk halvø inland ice uummannaq upernivik ø vaigat disko d g r disko bugt ilulissat aasiaat grønne ejland hareøen iti lli fau lt zo ne nordfjord mellemfjord ubekendt ejland uummannaq fjord qeqertarsuaq ? ? 50 km palaeogene intrusive complex lower palaeogene basalts maastrichtian–paleocene sediments albian–campanian sediments precambrian basement pre-volcanic fault fault with lateral or alternating displacements extensional fault ik kq qeqertarsuaq fig. 10. simplified geological map of the nuussuaq basin (after chalmers et al. 1999). ik, ikorfat fault zone; kq, kuu gannguaq–qunnilik fault; dgr, disko gneiss ridge. the offshore geology is indicated by paler shades. 21 sea-level curve, relative to present ma a ga td al fm eq al ul ik f m k an gi lia fm sl op e fa ul tco nt ro lle d sl op e un da te d se di m en ts k an gi lia fm it ill i f m u m iiv ik m b it ill i f m a na ri ar to rf ik m b vaigat fm kanisut mb maligât fm ifsorisoq mb unconformity associated with submarine canyon incision unconformity associated with valley incision unconformity tectonic phase uplift igneous activity lavas hyaloclastites deep marine sandstones and conglomerates fluviatile and deltaic sandstones and conglomerates deep marine thinly interbedded sandstones and mudstones mudstone coal chaotic beds turbidite channels ku ug an ng ua q– q un ni lik f au lt 1 2 65.5 61.1 58.7 55.8 70.6 83.5 85.8 88.6 93.6 99.6 112.0 albian cenomanian turonian coniacian santonian campanian maastrichtian danian selandian thanetian ypresianeoc pa le oc en e u pp er lo w er pa la eo ge ne c re ta ce ou s 22 k om e fm ba se m en t a llu vi al / la cu st ri ne / de lta ic a ta ne f m r av n k lø ft m b sl ib es te ns fje ld et f m a ta ne f m q ila ki ts oq m b a ta ne f m q ila ki ts oq m b it ill i f m a af fa rs ua q m b a ga td al fm eq al ul ik f m k an gi lia fm k an gi lia fm k an gi lia fm d el ta ic sl op e it ill i f m k us si ne ru ju k m b a ta ne f m k in gi tt oq m b a ta ne f m k in gi tt oq m b q ui ka vs ak fm q ui ka vs ak f m tu pa as at v al le y d el ta ic d el ta ic d el ta ic pa at uu tk lø ft en v al le y ? ? d is ta l s ub m ar in e lo be s it ill i f m u m iiv ik m b u pe rn iv ik n æ s fm un da te d se di m en ts es tu ar in e ba se m en t d ri ft sy nri ft po st -r ift sy nri ft 8 7 6 5 4 3 2 1 tss west greenland basalt group atanikerluk fm ? ? ? ? ? 3 4 5 8 7 6 8 1 2 3 4 56 7 fig. 11. tectonic events and regional depositional units, based on dam & nøhrhansen (2001). tss, tectonostratigraphic sequences; eoc, eocene. 1: gant#1 well; 2: gro#3 well and itilli valley area; 3: agatdalen; 4: ataata kuua and paatuut; 5: atanikerluk; 6: kussinerujuk; 7: slibestensfjeldet; 8: umiivik-1 well and itsaku. the chronostratigraphy and the sea-level curve were produced using tscreator pro v. 4.0.2 (2009; http://tscreator.org). compare with fig. 16. see text for further explanation. 23 the cretaceous–paleocene sedimentary succession of the nuussuaq basin onshore west greenland is best known from eastern disko and nuussuaq, with minor, and less well-known, outcrops in the northern part of the region on upernivik ø, qeqertarsuaq and svartenhuk halvø (fig. 10). seismic and other geophysical data indicate that the mesozoic succession is at least 6 km and possibly up to 10 km thick in the western part of the basin (christiansen et al. 1995; chalmers et al. 1999; marcussen et al. 2002). the eastern part appears to have much shallower depths to basement (chalmers et al. 1999), and this part of the basin could represent thermal subsidence following the initial rifting episode (chalmers et al. 1999). the outcrops record a complex history of rifting, subsidence and uplift commencing with an earliest cretaceous (or earlier) rift episode followed by a phase of thermal subsidence during the cenomanian – early campanian (fig. 11). rifting resumed in the early campanian and increased in the maastrichtian – early paleocene (dam & sønderholm 1998; dam et al. 2000; dam 2002), culminating during the early paleocene. the first phase of these later rift episodes was characterised by largescale normal faulting, whereas the later episodes were associated with continued extension and regional uplift (dam & sønderholm 1998; dam et al. 1998a, 2000; chalmers et al. 1999). the late phases were accompanied by widespread igneous activity and extrusion of a thick succession of flood basalts (fig. 12; a.k. pedersen et al. 2006a, and references therein). the exposed part of the succession in the nuussuaq basin can be divided into eight tectonostratigraphic sequences (tss; fig. 11); the early rift episode includes two sequences and the late episode six sequences. these sequences are mainly related to tectonic events marking discrete basin-fill phases (dam & nøhr-hansen 2001). tss 1. the oldest sediments exposed in the disko – nuussuaq basin represent a syn-rift episode of ?aptian– albian age represented by the kome and slibestensfjeldet formations (fig. 11). this rift episode is dominated by n–s extensional faults which, however, were also reactivated during later stages (l.m. larsen & pedersen 1990; chalmers et al. 1999). the n–s trend is expressed particularly by the disko gneiss ridge and this trend can be followed on western nuussuaq in the kuugan nguaq–qunnilik fault (figs 10, 12). the eastern boundary fault system has an overall nnw–sse trend but is segmented with individual segments trending n–s or nw–se (fig. 10; rosenkrantz & pulvertaft 1969; chalmers et al. 1999). the kome formation reflects an environment dominated by fluvial plains and local fan deltas amid basement highs. the kome for mation is overlain locally by lacustrine deposits of the slibestensfjeldet formation. tss 2. following the early rifting episode there was a long period of thermal subsidence that spanned the late albian/cenomanian – turonian – earliest campanian. it was initiated by a major flooding surface represented by offshore and deep marine deposits of the itilli formation to the north and west and by fluvio-deltaic and shallow marine deposits of the atane and upernivik næs formations to the east and south. the delta fanned out to the west and north-west from a point east of disko (figs 11, 12a; g.k. pedersen & pulvertaft 1992). on nuussuaq, the transition from shallow marine and fluviodeltaic deposition in the eastern part of the basin into deep marine deposition farther west was controlled by the n–s-trending kuugannguaq–qunnilik fault that crosses disko and nuussuaq (figs 10, 11, 12a). on svartenhuk halvø contemporaneous deep-water deposition in a slope setting is recorded by a thick distal turbidite succession assigned to the itilli formation (dam 1997). this unit includes marine anoxic shales of presumed cenomanian–turonian age, that are possibly the source for the marine itilli oil type (dam et al. 1998b; bojesen-koefoed et al. 1999). tss 3. in earliest campanian time a new tectonic episode was initiated that lasted from the early campanian to the paleocene (dam et al. 2000). the early phase of this rifting episode (tss 3) is represented by the aaffarsuaq member of the itilli formation and lasted into the maastrichtian. this phase is characterised by normal faulting, subsidence and syn-rift sedimentation. it resulted in the development of an angular unconformity, and deltaic deposition gave way to catastrophic deposition in a footwall fan setting along n–s-trending normal faults. in the eastern part of the region, uplift resulted in significant erosion of previously deposited atane formation deposits, and it is therefore expected that turbidite sandstone bodies of regional extent are present in the deep-water facies in the offshore basins to the west. facing page: fig. 12. palaeogeographic reconstructions of the nuussuaq basin during a: the cenomanian/turonian – earliest campanian (tss 2); b: the latest maastrichtian (tss 4); c: the danian (tss 5); d: the earliest selandian; e: early selandian volcanism, and f: early selandian dammed lake phase. see text for further explanation. 24 50 km disko nuussuaq a b c d e f svartenhuk halvø hinterland delta slope basin floor volcanoes fault lavas 25 26 tss 4–6. in late maastrichtian – early paleocene times, the stress system in the region changed and extension took place along nw–seand n–s-trending faults. these form the present eastern limit of the basin and displaced and rotated the major n–s-trending blocks in the basin (e.g. chalmers et al. 1999). this trend is identical to several shear zones in the precambrian basement east of disko bugt, suggesting that these shear zones exerted an influence on later faulting and the trend of a possible major transfer fault situated in the vaigat area (dam 2002, wilson et al. 2006). major faulting also occurred along the nw–se-trending faults and the rift blocks show evidence of major erosion before being covered by upper maastrichtian – lower paleocene marine sediments and middle paleocene volcanic rocks (e.g. dam & sønderholm 1998; dam et al. 1998a). birkelund (1965), rosenkrantz & pulvertaft (1969), j.m. hansen (1980b) and nøhrhansen (1996) noted that the cretaceous faunas and floras in the nuussuaq basin are similar to those of the north american interior seaway while there is an overwhelming european affinity in the danian, suggesting that an important change in palaeogeography and palaeoceanography took place during the latest cretaceous and earliest paleocene (rosenkrantz & pulvertaft 1969; j.m. hansen 1980b; nøhr-hansen & dam 1997). three major tectonic episodes have been recognised in the latest maastrichtian – earliest paleocene, each associated with incision of valley systems and development of submarine canyons. the first of these episodes (tss 4) is of latest maastrichtian age and is represented by the kangilia formation in which two major se–nw-trending submarine canyons have been documented from outcrops (figs 11, 12b). the second, earliest paleocene episode (tss 5) is represented by the tupaasat and nuuk qiterleq members of the quikavsak formation. it was associated with major uplift of the basin and fluvial valley incision into early paleocene fault scarps and was characterised by catastrophic deposition (figs 11, 12c). the third episode (tss 6) was associated with renewed uplift during the early paleocene, and valleys were incised into the old valley system (paatuutkløften member of the quikavsak formation). crossing the kuuganng uaq– qunnilik fault, the incised fluvial valleys pass westwards into a major submarine canyon system. the sand-dominated fill of this canyon system is referred to the marine agatdal formation, named after equivalent valley-fill sediments in central nuussuaq. this episode was followed by very rapid subsidence. the incised valleys were eventually filled with transgressive estuarine and shoreface deposits before they were blanketed by offshore tuffaceous mudstones referred to the eqalulik formation (figs 11, 12d) immediately prior to extrusion of picritic hyaloclastite breccias of the vaigat formation (figs 11, 12e). the recurrent episodes of uplift and incision of submarine canyons and valleys in atane formation deposits in the eastern outcrop area resulted in major redistribution of sandstones into the deep-water environments to the west, and major turbidite sandstone bodies are thus suspected to be regionally present. tss 7. extrusion of the volcanic succession can be divided into two phases and is related to continental break-up in the labrador sea region (a.k. pedersen et al. 2006a, and references therein). the first phase, of selandian to thanetian (late paleocene) age, was dominated by extrusion of olivine-rich basalts and picrites (figs 11, 12e) and later by more evolved, plagioclase-phyric basalts (vaigat and maligât formations of the west greenland basalt group). the first volcanism recorded in the nuussuaq basin took place in a marine environment and eruption centres were located in the westernmost part of the basin (fig. 12e). thick hyaloclastite fans of the anaanaa and naujánguit members prograded towards the east (figs 12f, 16). as the volcanic front moved eastwards, large lakes were formed between the volcanic front to the west and the cratonic crystalline basement to the east (fig. 12f), giving rise to synvolcanic lacustrine deposits (atanikerluk formation). tss 8. during the eocene, magmatic activity in the nuussuaq basin resumed with an episode of intrusion of dyke swarms and extrusion of basalts and sparse comendite tuffs of the kanísut member. the volcanic succession was dissected by n–s-trending faults and a new ne–sw fault trend (the itilli fault zone; figs 10, 11). the tectonic activity probably waned during late palaeogene time, and during the neogene the area was lifted by 1–2 km to its present elevation (chalmers 2000; bonow et al. 2007, and references therein; japsen et al. 2009). history. the nuussuaq group comprises the preand synvolcanic cretaceous–paleocene sediments on a number of islands and peninsulas in west greenland between 69° and 72°n (fig. 10). intrabasaltic sediments are not included in the nuussuaq group but in the overlying west greenland basalt group, with the local exception of thin, intra-volcanic wedges of sediment that demonstrably interdigitate with prograding hyaloclastite breccias and lavas of the lowermost west greenland basalt group. initial investigations of the geology date back to k.l. giesecke, who described the area in detail around 1810 (in: steenstrup 1910), rink (1853, 1857) and nordenskiöld (1871); the latter erected three litho stratigraphic units for the pre-volcanic sediment, separate from the intra-basaltic ifsorisok beds (fig. 13; see section on ‘previous work’ above). however, these units were not recognised by l. koch (1929) in his exhaustive account of the stratigraphy of greenland. he referred all the sediments to one formation – the nugsuak for mation – but also predicted that “future investigations will doubtlessly result in a division of the beds which i have included in one formation, into several formations” (l. koch 1929 p. 258). a more detailed lithostratigraphical scheme was developed as a result of the work during the nûgssuaq expeditions 1938–1968 (see section on ‘previous work’ above; rosenkrantz 1970). this was, however, rather incomplete and included units, especially at member level, that were not formally described (fig. 13). name. after the nuussuaq peninsula, where the most extensive outcrops occur (fig. 10). type area. the south coast of nuussuaq, where the most complete and best exposed sections of the nuussuaq group occur, e.g. at atanikerluk, the coastal slopes at paatuut, along the western slope of the ataata kuua river, and in river gorges in the southern part of the itilli valnuussuaq group new group nordenskiöld 1871 troelsen 1956 rosenkrantz 1970 henderson et al. 1976 present paper sinnifiklagren* ifsorisoklagren* komelagren kome fm kome fm maligât fm maligât fm vaigat fm kome fm slibestensfjeldet fm ifsorisok fm* ifsorisok fm* ifsorisok mb* nugsuak formation koch 1929 patoot fm pautut fm öfre atanekerdluk-lagren atanelagren (n. atanekerdluklagren) upper atanikerdluk fm u. atanikerdluk fm atanikerluk fm eqalulik fm agatdal fm agatdal fm agatdal fm vaigat fm kangilia fmkangilia fm kangilia fm atane fm n uu ss ua q g ro up w g bg atane fm atane fm upernivik næs fm u pe rn iv ik n æ s fm itilli fm marine upper cretaceous marine upper cretaceous sinnifik fm* quikavsak fm * intrabasaltic sediments fig. 13. comparison of former lithostratigraphical subdivisions and the scheme used in this paper. all pre-volcanic and the earliest synvolcanic sedimentary rocks of the nuussuaq basin constitute the nuussuaq group (new), which comprises 10 formations (new or revised). wgbg, west greenland basalt group. 27 quikavsak formation atane formation atane formation kangilia formation kangilia formation vaigat formation d b d d eqalulik formation quikavsak formationvaigat formation quaternary cover south north kangilia formation atane formation qilakitsoq member streamintrusion atanikerluk formation 300 200 200 m 500 400 600 700 m above sea level kangilia formation, very poorly exposed fig. 14. the western side of the ataata kuua river gorge, in the type area of the nuussuaq group; d, dolerite dyke, b, burnt mudstones. for location, see fig. 40; for scale, see fig. 15. fig. 15. the western side of ataata kuua, in the type area of the nuussuaq group, illustrates two phases of incision. the marine mudstone and turbidite sandstones of the kangilia formation represent a late maastrichtian submarine canyon incised into the santonian deltaic atane formation; mudstones of the kangilia formation are shown in brown, sandstones and conglomerates in beige. the danian quikavsak formation represents a fluvial system incised into the kangilia formation. photogrammetrically measured section, from a.k. pedersen et al. (2007b). borehole ggu 247801 is located c. 600 m north of the outcrop shown here (figs 40, 43). 28 ley (figs 2, 40, 65). the section along the west slope of ataata kuua has been measured photogrammetrically (figs 14, 15; a.k. pedersen et al. 2007b). distribution. the nuussuaq group outcrops in west greenland between 69° and 72°n on disko, nuussuaq, upernivik ø, qeqertarsuaq, itsaku and svartenhuk halvø and has also been recorded in a number of shallow and deep wells in the region (figs 2, 16, 17). possible palae ogene sediments have previously been reported on angiissat, the south-easternmost island of a group of small islands in disko bugt named grønne ejland (henderson et al. 1976 p. 345). differentiation of the albian–cenomanian kome, upernivik næs and atane formations that form the lower part of the group can be difficult because these formations all include marginal marine deposits that are dominated by sandstones and have some sedimentary facies in common. for practical purposes, the kome formation is restricted to areas in northern nuussuaq, east of ikorfat, where alluvial sediments are in contact with basement. the upernivik næs formation, which occurs north of nuussuaq, comprises marginal marine sediments, while the fluviodeltaic atane formation is restricted to the disko–nuussuaq area. thickness. composite sections of outcrops and wells show thicknesses of up to c. 3 km for the nuussuaq group (fig. 17). in the western part of the area, seismic and magnetic data suggest that the mesozoic sediments are at least 6 km and possibly as much as 10 km thick (chri stiansen et al. 1995; chalmers et al. 1999). a seismic section across the vaigat indicates that the non-marine cretaceous section is at least 2500–3000 m thick (fig. 44; marcussen et al. 2002). lithology and depositional environment. the exposed part of the nuussuaq group consists entirely of siliciclastic sediments (figs 16, 17). a reconnaissance study of the petrology of the cretaceous and paleocene sandstones revealed varying feldspar contents and three types of cement: carbonates, silica and fe-hydroxides. based on differences in texture, schiener (1975) suggested two provenance areas: the archaean crystalline basement rocks (angular feldspar grains) and older sedimentary rocks that have been recycled (well-rounded quartz grains). detrital zircon dating of the sediments indicates that most of the cretaceous sediments were transported from areas with archaean basement, whereas the late cretaceous and paleocene sediments at itsaku on svartenhuk halvø contain proterozoic zircons, probably derived from the prøven igneous complex (scherstén & sønderholm 2007). the basal lower cretaceous part of the succession includes coarse-grained, syn-rift breccias and conglomerates that onlap basement highs and are overlain by sandstones, heteroliths and mudstones of non-marine, mostly alluvial origin (fig. 16). slightly younger braided river sandstones interbedded with sandstones, heteroliths and mudstones deposited in tidal estuarine and coastal plain environments crop out on upernivik ø and qeqertarsuaq and are referred to the upernivik næs formation. on northern nuussuaq, the kome formation is overlain by lacustrine mudstones and sandstones of the slibestensfjeldet formation. the atane formation un conformably overlies the slibestensfjeldet formation and possibly also the kome formation. the atane formation is dominated by coarsening-upward successions of mudstones and sandstones deposited in a deltaic and fluvial environment during the albian to santonian. across the kuugannguaq–qunnilik fault (fig. 10), the deltaic sednext pages 32-33: fig. 16. stratigraphy and depositional settings of the formations and members of the nuussuaq group. the sections are located in fig. 2 or in the detailed maps located in fig. 2. up., uparuaq qusuitsut (svartenhuk halvø). formations are shown on the left hand side of the logs and members on the right hand side. the vertical axis of the diagram indicates the approximate age of the lithostratigraphical units (data on ages are found in the selected numbered references below). compare with fig. 17. stratigraphic abbreviations: a, assoq member (atanikerluk formation); ak, akunneq member (atanikerluk formation); an, anaanaa member (vaigat formation); at, atanikerluk formation; ac, annertuneq conglomerate member (kangilia formation); e, eqalulik formation; it, itivnera bed (atane formation); m, maligât formation; n, naujât member (atanikerluk formation); na, naujánguit member (vaigat formation); nq, nuuk qiterleq member (quivaksak formation); o, ordlingassoq member (vaigat formation); oac, oyster–ammonite conglomerate bed (kangilia formation); p, pingu member (atanikerluk formation); pa, paatuutkløften member (quikavsak formation); rdm, rinks dal member (maligât formation); sv, svartenhuk formation; tu, tupaasat member (quikavsak formation); u, umiussat member (atanikerluk for mation). references: 1, birkelund (1965); 2, boyd (1998 a); 3, christiansen et al. (1999); 4, christiansen et al. (2000); 5, dam & nøhr-hansen (2001); 6, dam et al. (1998b); 7, dam et al. (1998c); 8, dam et al. (2000); 9, koch (1964); 10, koppelhus & pedersen (1993); 11, lanstorp (1999); 12, j.g. larsen & pulvertaft (2000); 13, nøhrhansen (1996); 14, nøhr-hansen et al. (2002); 15, olsen & pedersen (1991); 16, a.k. pedersen (1985); 17, a.k. pedersen et al. (2006b); 18, piasecki et al. (1992); 19, storey et al. (1998). 29 thanetian svartenhuk halvø western nuussuaq central nuussuaq aaffarsuaq e of ik or fa t u m iiv ik it sa ku u p. w o f i ko rf at north coast nuussuaq upernivik næs, qeqertarsuaq selandian pa le oc en e u pp er c re ta ce ou s lo w er c re ta ce ou s danian maastrichtian campanian santonian turonian cenomanian albian coniacian va ig at k an gi lia va ig at k an gi lia it ill i u pe rn iv ik n æ s u pe rn iv ik n æ s sl ib es te ns fje ld et a ta ne k om e u m iiv ik u m iiv ik 12 17 19 17 4 15 1 6 1 7 4 12 4 9 2 3 sv ? ravn kløft kingittoq contact to basement contact to basement it ill i it ill i a ta ne k an gi lia k an gi lia it ill i va ig at a ga td al ac e e va ig at e na o 14 na an o na it o a af fa rs ua q q ila ki ts oq a na ri ar to rf ik 1.5 km undated deep marine sediments 8 13 3 t tuff tt t volcanic rocks coal mudstone sandstone conglomerate macrofossils key reference hyaloclastic breccia subaerial lava flow 30 central nuussuaq south coast nuussuaq south coast nuussuaq north coast disko eastern disko south coast disko agatdalen ataata kuua paatuut kingittoq atanikerluk asuk pingu skansen, assoq 13 14 17 5 contact to basement a ta ne a ta ne a ta ne a ta ne sk an se n a ta ne sk an se n it ill i k in gi tt oq k in gi tt oq k us si ne ru ju k va ig at va ig at va ig at a ga td al k an gi lia q ui ka vs ak q ui ka vs ak k an gi lia it ill i e e m na at oac o pa nq tu o n rdm a u a au p ak n tu rdmrdm o na rdm rdm a af fa rs ua q q ila ki ts oq q ila ki ts oq a ta ne 8 13 2 15 16 11 10 10 mm at 17 16 19 17 18 m at m at t t t alluvial fan fluvial channel flood plain lake delta front marine shoreface marine shelf marine slope marine slope channel delta plain estuary 31 va ig at va ig at o rd lin ga ss oq n au ja ng ui t q ila ki ts oq n au ja ng ui t q ila ki ts oq va ig at va ig at a ta ne a ga td al a na an aa k an gi lia k an gi lia it ill i k ak ili sa at n er ut us oq u pe rn iv ik n æ s u pe rn iv ik n æ s it ill i e e na o e e ag k i va ig at it ill i it ill i a ta ne svartenhuk halvø gro#3 gank#1 kangiliaitsaku umiivik agatdalen tunoqqu aaffarsuaqupernivik næs qeqertarsuaq 7 11 9 813 12 15 14 500 m k fig. 17. simplified logs showing the distribution and thickness of the formations of the nuussuaq group. members are only indicated for the atane formation and for the volcanic formations. in most areas, the lower boundary of the nuussuaq group is not seen, and the thicknesses indicated for the formations are therefore minimum values. note that the marginal marine deposits (kome, slibestensfjeldet, upernivik næs and atane formations) are thick on disko, eastern nuussuaq and northwards to qeqertarsuaq. deep marine deposits (itilli, kangilia and agatdal formations) are thickest on western and northern nuussuaq and on svartenhuk halvø. at localities 1, 3, 5 and 6, the mudstones of the atanikerluk formation are interbedded with volcaniclastic breccias or subaqueous lava flows that have chemical compositions characteristic of the vaigat and maligât formations. na, naujánguit member; o, ordlingassoq member; rdm, rinks dal member; wgbg, west greenland basalt group. compare with fig. 16. 32 kome formation quikavsak formation (q) atane formation atanikerluk formation (at) 1 1 2 upernivik næs formation eqalulik formation (e) itilli formation (i) vaigat formation maligât formation agatdal formation (ag) slibestensfjeldet formation kangilia formation (k) wgbg n au ja ng ui t o rd lin ga ss oq n au ja ng ui t q ila ki ts oq o rd lin ga ss oq q ila ki ts oq o e e ii at at q k q o rd lin ga ss oq va ig at va ig at a ta ne a ta ne va ig at va ig at a ta ne sl ib es te ns fj. k om e k in gi tt oq r av n k lø ft k in gi tt oq k in gi tt oq r in ks d al m al ig ât a ta ni ke rl uk a ta ne m al ig ât a ta ni ke rl uk a ta ne sk an se n r d m a ta ne a ta ni ke rl uk m al ig ât sk an se n r d m q a ta ne at ataata kuua paatuut qorlortorssuaq asuk pingu skansen assoq kingittoq atanikerluk 6 5 slibestensfjeldet 10 4 2 13 upernivik næs71° 72° 70° 69° qeqertarsuaqsvartenhuk halvø disko 50 km nuussuaq 53° 3 4 56 7 8 9 1011 12 13 14 15 palaeogene volcanics precambrian basement cretaceous–paleocene sediments 33 iments of the atane formation pass westwards into marine, fault-controlled slope deposits. these sediments are referred to the itilli formation, of cenomanian to maastrichtian age, and comprise mainly homogeneous mudstones, thinly interbedded mudstones and sandstones, and turbidite channel sandstones. in the aaffarsuaq valley, an angular unconformity separates the deltaic atane formation from campanian turbidite deposits of the itilli formation (dam et al. 2000). major uplift during the maastrichtian resulted in incision of underlying units; submarine canyon incision is seen at several localities on nuussuaq, viz. at ataata kuua, at kangilia, in agatdalen and west of the kuuganng uaq–qunnilik fault (fig. 15, see plate 3). at these localities, the valley and canyon fill consists of successions of homogeneous sandstones and conglomerates, in many instances resulting from catastrophic gravity flow deposition. intercalated mudstones represent deposition from turbidity currents. these deposits are referred to the kangilia formation. renewed uplift during the danian generated new incision and fluvial to tidal valley deposits separated by lacustrine sediments filled the fluvial valleys. the lacustrine deposits comprise thin coarseningupward successions composed of mudstones and sand stones with in situ and drifted tree trunks. these sed i ments are referred to the quikavsak formation (dam 2002). the pre-volcanic phase culminated with major subsidence of the basin and deposition of sandstones from sediment gravity flows (the agatdal formation). during this period, the disko area was characterised by sedimentary bypass. the next phase of nuussuaq group sedimentation coincided with submarine volcanism west of nuussuaq. the oldest rocks of the west greenland basalt group are hyaloclastic breccias of the vaigat formation (hald & pedersen 1975; a.k. pedersen 1985; a.k. pedersen et al. 1993, 1996, 2002, 2006b; g.k. pedersen et al. 1998). the contemporaneous siliciclastic marine sediments com prise mudstones, thinly interbedded sandstones and mudstones and chaotic beds. volcaniclastic sandstones and conglomerates deposited from turbidity currents are common in this part of the succession. the marine synslibestensfjeldet formation atane formation itilli formation e vaigat formation precambrian kome formation fig.18. the sediments onlapping and overlying the basement high at ikorfat on the north coast of nuussuaq (view towards the south) represent the base of the nuussuaq group. the basal sediments are represented by the kome formation overlain by the slibestensfjeldet and atane formations. the highest peak is c. 2000 m a.s.l.; for location, see fig. 22. e, eqalulik formation. 34 volcanic sediments are referred to the eqalulik formation from the latest danian (nøhr-hansen et al. 2002). as the early paleocene volcanic breccias of the vaigat for mation (west greenland basalt group) prograded eastwards from sources west of the present coastline, the formerly marine basin was dammed and large lakes formed on eastern nuussuaq and disko. the lakes were filled from the west by hyaloclastite breccias while siliciclastic sedimentation continued from the south-east. these sediments are re ferred to the atanikerluk formation and consist of shales with thin tuff beds and fine-grained sandstones deposited in two coarsening-upward successions. fossils. a rich flora and fauna from the group has been described. the fauna numbers several hundred species including bivalves, gastropods, nautiloids, ammonites, belemnites, echinoderms, cirripeds, brachiopods, corals, decapod crustaceans, serpulids, fish, ostracods, foraminifera, bryozoans and insects. fossil invertebrates are best known from the itilli, kangilia and agatdal formations. the rich flora comprises more than 600 species of plant leaves, wood, fruits, spores, pollen and dinoflagellate cysts (hereafter named dinocysts). well-preserved plant fossils are mainly found in the kome, atane, quikavsak, agatdalen and atanikerluk formations (for summary of studies, see section on ‘previous work’ above). dinocysts form the basis of biostratigraphy and correlation in recent studies. boundaries. the lower boundary of the group is only exposed on the north coast of nuussuaq at kuuk, vest erfjeld, talerua and ikorfat (fig. 18), on the east coast of itsaku and on svartenhuk halvø where the group overlies metamorphic basement that locally is strongly weathered. the boundary was also penetrated in the fp93-3-1 borehole on northern disko where the atane formation overlies a basement high (fig. 19). the upper boundary with the west greenland basalt group is diachronous (hald & pedersen 1975; a.k. pedersen et al. 1993). during the paleocene, volcanic breccias of the vaigat formation prograded eastwards from offshore sources and isolated the formerly marine basin in eastern nuussuaq and disko, eventually onlapping the precambrian basement east of the ikorfat fault zone (l.m. larsen & pedersen 1992; a.k. pedersen et al. 1996, 2002; g.k. pedersen et al. 1998). geological age. dating of the exposed and drilled parts of the basin is mainly based on ammonites, corals, foraminifera, coccoliths, spores, pollen, dinocysts and macroflora fossils. the fossils indicate that outcrops of the nuussuaq group span the albian to the upper danian – selandian. brackish-water dinocysts from sediments onlapping the basement indicate a pre-late albian age, and this is supported by the scarcity of angiosperms clay silt sand pebbles vf f m c vc f m c 130 110 120 100 90 80 60 70 50 40 30 m fp93-3-1 q ua te rn ar y a ta ne f or m at io n pr ec am br ia n gn ei ss strongly weathered surface fig. 19. detailed sedimentological log of the atane formation overlying weathered gneiss basement in the the fp93-3-1 core in the kuugannguaq valley on north-west disko. the well dips 60° and driller’s depths are shown. the stratigraphic thickness of the atane formation is c. 50 m; the sediments are not referred to a member. for location, see fig. 2. depositional environments are indicated by colours in the left column of the log (see plate 1). 35 36 in the ikorfat flora (boyd 1998a, b, c, 2000). in the umiivik-1 well, the oldest dateable marine dinocysts are of turonian age; these were recorded 800 m above the base of the well (dam et al. 1998b). in the gro#3 well, the oldest dateable dinocysts are of coniacian age and were recorded 1500 m above the base (christiansen et al. 1999). consequently, the deeper parts of these wells may have penetrated successions older than albian. age-specific biomarker data from oil seeps at several outcrops of the nuussuaq group suggest that source rocks of aptian–albian and cenomanian–turonian age are present in the nuussuaq basin (bojesen-koefoed et al. 2004, 2007). 40ar/39ar dating of the volcanic rocks coeval with the upper levels of the nuussuaq group (the vaigat formation and the rinks dal member of the maligât formation) indicates an age of 60.4 ± 0.5 ma for these rocks (storey et al. 1998). this age has been recalculated to 60.8 ± 0.5 ma based on recalibration of the fish canyon tuff standard (skaarup & pulvertaft 2007). correlation. the nuussuaq group is coeval with the cre taceous hassel and kanguk formations and the uppermost cretaceous–oligocene eureka sound group of the canadian arctic islands (e.g. miall et al. 1980; miall 1986; burden & langille 1990; harrison et al. 1999). it constitutes an onshore analogue to the cretace ous–paleocene basins offshore west greenland (rolle 1985; chalmers & pulvertaft 2001; dalhoff et al. 2003; sørensen 2006; gregersen et al. 2007) and to the cretaceous bjarni, markland and cartwright formations offshore eastern canada in the davis strait and labrador sea (balkwill et al. 1990; government of newfoundland and labrador 2000; sønderholm et al. 2003). subdivision. the nuussuaq group is subdivided into the kome, slibestensfjeldet, upernivik næs, atane, itilli, kangilia, quikavsak, agatdal, eqalulik and atanikerluk formations (figs 13, 17). kome formation redefined formation history. the geology of the kuuk area was described by giesecke in 1811 (diary entry for 18 june; in: steenstrup 1910). he mentioned that the coal seams provided fuel for ‘kolonien omenak’, now the town of uummannaq (fig. 20). a 0.7−1.3 m thick coal seam was exposed and exploited until 1832 (rink 1857). the mineralogical museum in copenhagen received 21 samples of “shale with remains and imprints of different fossil plants, especially ferns” collected near kome (rink 1855 p. 214). heer (1883b) described an early cretaceous fossil flora from the north coast of nuussuaq that he termed the kome flora. according to nordenskiöld (1871), several of these fossils had been collected by rink and were sent to heer from the mineralogical museum in copenhagen. the kome formation was defined in the area between kuuk and ikorfat by nordenskiöld (1871 p. 1040) and described as follows (translated by the authors): “the komelagren [kome formation] constitutes the older part of the cretaceous, according to heer (1868). the name designates a sedimentary, coal-bearing formation exposed at various localities between kuuk and ikorfat along the coast of nuussuaq, south-west of uummannaq. the name originates from the outcrop which has the most important coal bed and where giesecke and rink most likely collected plant fossils. these beds [komelag ren, the kome formation] are not restricted to kuuk, but are found, except where interrupted by gneiss highs, along the entire coastal section between kuuk and ikorfat” (figs 2, 21, 22). nordenskiöld (1871) proposed a stratigraphy with a geographically restricted lower cretaceous kome formation overlain by the upper cretaceous atane formation which also covered southern nuussuaq and northern disko. gry (1940) interpreted the boundary between the kome and the atane formations as an un conformity. koch (1964) described the kome formation as occurring almost continuously between kuuk and ikorfat, interrupted only between angiarsuit and ujarattoorsuaq where younger upper cretaceous deposits are downfaulted (figs 21, 22). he suggested that the beds that unconformably overlie the kome formation on the north coast of nuussuaq could be equivalent to the atane for mation or, alternatively, to the upernivik næs formation. schiener (1977) did not find convincing evidence for an unconformity within the section at ikorfat and concluded that the majority of the sediments were deposited during the early cretaceous. he implied that the atane formation is absent on northern nuussuaq east of the ikorfat fault zone (fig. 10). in the kuuk area, the kome formation was interpreted to comprise fluviodeltaic deposits (pulvertaft 1979) that were suggested to form part of the atane formation by g.k. pedersen & pulvertaft (1992). in contrast, midtgaard (1996b) distinguished five sedimentary units in the area between ikorfat and qaarsut and the lowest of these is here assigned to the kome formation. name. in his descriptions of the coal seams on northern nuussuaq that provided fuel for ‘kolonien omenak’ (the town of uummannaq), giesecke in 1811(in: steenstrup 1910) spelled the name either ‘koome’ or ‘kook’. thalbitzer (1910) explained ‘koome’ as casus locativus of ‘kook’, e.g. ‘in or at kook’. the spelling ‘kome’ was used by rink (1857) and heer (1882, 1883b). rink (1853 p. 175) wrote: “the settlement of kome lies at the mouth of the large valley tuëparsoït [now tua passuit, fig. 21]. between this and sarfarfik a broad open valley is found. this contains a small river (kook) thus the name of the site, where it reaches the shore” (translation from danish by the authors). the site is shown as kook, kûk or kuuk on newer maps (fig. 21). distribution. the kome formation is presently known from the coastal area on the north coast of nuussuaq between kuuk and ikorfat (figs 21, 22). gry (1940, 1942) briefly described sediments onlapping the basement on itsaku (fig. 73) and referred these to the kome formation on the basis of plant fossils from the lower part of the sedimentary succession. these sediments were tentatively referred to the atane formation by chri stiansen et al. (2000) and j.g. larsen & pulvertaft (2000) on the basis of sedimentological similarities. in the present paper they are referred to the upernivik næs formation. east of saqqaqdalen on southern nuussuaq (fig. 2), outcrops of fluvial pebbly sandstones overlying the basement are preserved in a small down-faulted area (pulvertaft 1989a, b) and are tentatively assigned to the kome formation. type section. the type section is at majorallattarfik, immediately west of the kuuk delta (figs 21, 23). the base of the type section is located at 70°38.60´n, 52°21.83´w. reference sections. reference sections have been measured at slibestensfjeldet and ikorfat (midtgaard 1996b) (fig. 22). these sections illustrate the marked lateral and stratigraphic, lithological changes in the formation from stacked conglomerates to coarsening-upward, fine-grained sandstones and mudstones (figs 24, 25). thickness. the thickness of the kome formation varies from c. 25 m at vesterfjeld to more than 150 m between talerua and ikorfat, and 140 m at majorallattarfik near kuuk (figs 21, 22). lithology. the basal part of the kome formation overlies and onlaps precambrian basement at kuuk, talerua and ikorfat. the basement rocks are in places weathered to a depth of 35 m (heim 1910; gry 1942; pulvertaft 1979; midtgaard 1996b). the uppermost part of the basement consists almost exclusively of quartz grains within a powdery matrix of kaolinite. the gneiss has a greenish colour probably due to alteration of biotite to chlorite (heim 1910; pulvertaft 1979). the basal sediments are typically one of three main facies: (1) poorly sorted sandstones, rich in kaolinite and devoid of sedimentary structures, (2) diamictites, with blocks of vein quartz in a sandy clay matrix, or (3) unsorted conglomerates with poorly rounded quartz boulders and slabs of silty mudstones (fig. 26; gry 1942; schiener 1977; pulvertaft 1979). at talerua (between vesterfjeld and fig. 20. view from uummannaq towards the coal-bearing deposits of the kome formation at kuuk on the north coast of nuussuaq. the coal was being mined as early as the 18th century for use in uum mannaq. view foreshortened due to the use of a telescopic lens. for location, see figs 2, 21. 37 ikorfat), the basal sediments are breccias with large clasts comprising both gneiss with weathered feldspars and carbonaceous mudstones (fig. 25a). at ikorfat, breccias with angular gneiss clasts are overlain by a conglomerate bed, a few metres thick, comprising subangular clasts of quartz and intensely kaolinised feldspars. the diamictites and poorly sorted conglomerates are not seen above the basement highs. the poorly sorted, coarsegrained deposits are overlain by thin conglomerates, cross-bedded, locally channellised, coarse-grained sandstones, mudstones and thin discontinuous coal beds (fig. 25). root horizons are frequent and the mudstones commonly contain abundant comminuted plant debris (fig. 25; schiener 1977; pulvertaft 1979; midtgaard 1996b). fine-grained sandstones with wave-ripples forming coarsening-upward successions occur in the upper part of the formation and are common in the ikorfat area. the topmost bed of the kome formation is a bleached, crumbly mudstone with abundant root-casts overlain by a coal bed (fig. 25c; midtgaard 1996b). 2 km 70°40′ 70°35′ 52°25′ 200 400 200 800 800 1000 1600 10 00 1000 14 00 12 00 400 600 600 15 15 14 11 17 8 majorallattarfik tu ap as su it sarfâgfik kûksa rfa aff iip k uu ssi ne rsu a kuuts iaq ku uk vaigat formation tunoqqu member vaigat formation kome formation precambrian ice sea/lake fault abandoned settlement inferred base of cretaceous sediments quaternary cover fig. 21. geological map of the area around kuuk (from pulvertaft 1979). majorallat tarfik is the type locality of the kome formation. the sediments overlie weathered basement and are well exposed. for location, see fig. 2. 38 in the kuuk area, the sediments overlying the coarsegrained, basal deposits consist of three facies associations: a, b and c (figs 24, 27, 28). association a is dominant and consists mainly of trough cross-bedded, erosionally based sandstones in shoe-string or irregular tabular bodies up to 5 m thick (figs 24, 28). the sandstones are coarse-grained, often gravelly, with angular clasts, and is poorly sorted. it contains slightly to moderately kaolinised feldspars and kaolinite cement. fragments of coalified wood and fine-grained plant debris are common. the orientation of shoe-string bodies and dip direction of foresets indicates transport from the south-east. the sandstones alternate with dark laminated silty mudstones containing much coalified plant debris, thin lenses of lustrous, coalified, woody tissue and thin coal seams (fig. 24; pulvertaft 1979). association b consists of distinctive coarsening-upward successions of mudstones, cross-laminated siltstones, sandstones and locally thin layers of gravelly sandstones (figs 24, 27). association c is dominated by tabular cross-bedded, medium-grained, moderately well-sorted subarkosic sandstones with pebble lags of quartz or chert. comminuted coalified plant debris and mica are common in the sandstones. foresets dips are to the north-west. cross-laminated, fine-grained 53°00’ 70°45’ 52°45’ 2 km qaarsut talerua airportikorfatikorfat qaqqarsuiqaqqarsui vesterfjeldvesterfjeld kus s in er su a ikorfat d 2525 30 30 20 28 b c a q aa rs u t k us s i ne rs ua t kus s in er su a slibestensfjeldet vesterfjeld ikorfat qaqqarsui østerfjeld aarru sa p ku ss in ika ss aq an gia rs ui t n ip ita rt uk as sa at u jar at to or su up k illi a u jar at to or su aq u jar at to or su up k an gil ia sio ra rt ar ta rfi k j.p.j. ravn kløft intrusion eqalulik formation atane formation kingittoq member atane formation ravn kløft member slibestensfjeldet formation atane formation undivided, west of ikorfat itilli formation umiivik member kome formation precambrian ice sea/lake fault strike and dip of strata vaigat formation tunoqqu member vaigat formation quaternary cover fig. 22. detailed geological map of the north coast of nuussuaq, from qaarsut to ikorfat. the locations of the four sedimentological logs (a–d) in fig. 25 are indicated. contour interval 200 m. modified from rosenkrantz et al. (1974) and midtgaard (1996b). for location, see fig. 2. 39 sandstones overlying the medi um-grained sandstones form a minor part of the association. wave-generated sedimentary structures or mud-draped foresets indicating tidal influence have not been seen (pulvertaft 1979). facies association c occurs at the top of the sections at kuuk and majorallattarfik (fig. 28; t.c.r. pulvertaft, personal communication 2008). fossils. the kome formation contains macrofossil plants referred to the kome flora (heer 1883a). the genera in the kome flora were revised by k.r. pedersen (1976). a separate ikorfat flora has been distinguished by boyd (1998a, b, c, 2000) in the ikorfat area and is suggested to be slightly older than or contemporaneous with the kome flora (boyd 1998a, b, c, 2000). samples of mudstones collected by croxton (1978a, b, section c20) at majorallattarfik near the top of the section at kuuk contain a poor assemblage of brackishwater dinocysts, spores and pollen. few specimens of the dinocyst vesperopsis nebulosa occur in the lower part, whereas few specimens of nyktericysta davisii and pseu doceratium interiorense occur in the upper part. the presence of the spore cicatricososporites auritus may indicate a middle to late albian age, whereas the absence of the pollen rugubivesiculites rugosus indicates a pre-late albian age. the presence of nyktericysta davisii may indicate a younger age, as at other localities this species first occurs in the overlying ravn kløft member (atane formation). samples of mudstones collected by midtgaard (1996b) from the upper part of the kome formation reference section at slibestensfjeldet (fig. 25c) also contain ves peropsis nebulosa and common pseudoceratium interiorense and, in common with the kuuk section, the pollen rugubivesiculites rugosus is absent. assemblages dominated by pseudoceratium interiorense are known from the slibestensfjeldet formation, whereas assemblages dominated by nyktericysta davisii first occur in the ravn kløft member of the atane formation between vesterfjeld and ikorfat (fig. 29). croxton (1978a, b) noted the absence of angiosperm pollen and the presence of spores such as cicatricosisporites, gleicheniidites, pilosisporites and vitreisporites, indicating an early to midcretaceous age. depositional environment. the diamictites and poorly sorted conglomerates in close association with the basement highs are interpreted as mass-flow deposits (figs 25, 29); talus cones and alluvial fans filled the topographic lows on the basement surface with immature sediments (midtgaard 1996b). the lack of mass-flow deposits above the basement highs suggests that the ini80 90 100 110 m 70 60 k om e fo rm at io n pr ec am br ia n ba se m en t 50 40 30 20 10 0 clay silt vf f m c vc sand fig. 23. type section of the kome formation at majorallattarfik, immediately west of kuuk, redrawn from croxton (1978a, b). the uppermost c. 20 m of the formation are poorly exposed and are not shown on the log. note that the facies illustrated in detail in fig. 24 recur here in a thicker succession. for location, see fig. 21; for a possible correlation between kuuk and slibestensfjeldet, see fig. 29; for legend, see plate 1. 40 0 1 2 3 4 5 6 7 8 9 10 11 12 13 14 m 0 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 m mud sand gravel mud sand gravel sandstone/pebbly sandstone, structureless laminated sandstone cross-bedded sandstone lithology and structures depositional environment ripple cross-laminated sandstone convolute lamination burrows coalified wood fragments coal laminated mudstone/siltstone weakly laminated mudstone fluvial channel flood plain facies association a of pulvertaft (1979) facies association bdelta front 41 fig. 24. reference section of the kome formation at kuuk. the log to the left is representative of facies association a, while association b is seen from 3.1–9.3 m in the log to the right (from pulvertaft 1979). for location, see fig. 21; for legend, see plate 1. 42 0 sl ib es te ns fje ld et f m k om e fo rm at io n 0 10 20 30 40 m 10 20 30 38 mud 242 m a.s.l. sand gravel 0 10 20 30 36 m m mud 2 m a.s.l. sand gravelmud section base 20 m above sea level (a.s.l.) sand gravel ★ ★ a b c k om e fo rm at io n k om e fo rm at io n fig.25. reference sections of the kome formations illustrating abrupt facies change and numerous horizons with rootlets. a and b are from the lower part of the formation and c and d are from the uppermost part of the formation. sections c and d also show the overall decrease in grain-size towards the west. the boundary between the kome formation and the overlying slibestensfjeldet formation is shown in section c. from midtgaard (1996b). for legend, see plate 1; for location, see fig. 22. 43 tial relief was levelled out as sediments accumulated and buried the highs. the alluvial fans were succeeded by fan deltas representing both subaerial and subaqueous deposition; the latter increases towards ikorfat (fig. 25d). the sparse dinocyst assemblage indicates that the water was intermittently brackish. the coarse sandstones of facies association a at kuuk are interpreted as fluvial deposits. their low textural and mineralogical maturity together with palaeocurrent indicators point to derivation directly from a weathered basement to the south-east, without reworking en route, and to rapid sedimentation. the dark mudstones are interpreted as overbank deposits of a floodplain environment or deposition in shallow lakes. the mudstones of facies association b are interpreted as having been deposited from suspension in bodies of standing water, presumably interdistributary bays or lakes (fig. 23, 24). sedimentary structures in the sandstones of facies association c suggest deposition from a sandy braided river flowing from the south, and the higher textural and mineralogical maturity of the sandstones indicates more protracted fluvial transport (pulvertaft 1979). the palaeosol that caps the kome formation represents a widespread regressive event (figs 25c, 30; midtgaard 1996b). boundaries. the kome formation overlies an undulating precambrian gneiss topography, filling valleys or depressions (henderson et al. 1976; schiener 1977; pulvertaft 1979; midtgaard 1996b). the lower boundary of the formation is exposed at kuuk, ikorfat and talerua (figs 2, 21, 22). in the coastal cliffs of northern nuussuaq, the relief on the gneiss surface can be seen to exceed 225 m (fig. 22). at vesterfjeld, the basement is weathered to a depth of 30 m with a gradual downward transition into unaltered gneiss. in western disko (stordal), the precambrian basement is exposed, locally with a 5−10 m thick weathered cap (a.k. pedersen & ulff-møller 1980). in other exposures, the gneiss is relatively unweathered (schiener 1977; midtgaard 1996b). the lower boundary of the kome formation thus displays a variety of basement−sediment transitions. the upper boundary of the kome formation is a sharp lithological boundary that varies very little laterally (figs 25c, 30). it is placed on top of a palaeosol horizon and was interpreted as a sequence boundary by midtgaard (1996b). previous workers, e.g. nordenskiöld (1871), gry (1942) and midtgaard (1996b), have considered the kome formation to be overlain by the atane formation on nuussuaq. we propose that the new slibestensfjeldet formation separates the kome and atane 0 10 20 30 40 mud 290 m a.s.l. sand gravel m d k om e fo rm at io n 44 formations in the area between ikorfat and slibe stensfjeldet. geological age. various fossils contribute to the determination of the geological age of the kome formation. plant macroflora suggests a barremian to aptian age (heer 1882; k.r. pedersen 1968). the ikorfat flora suggests an early to middle albian age (boyd 1998a, b, c, 2000). based on spores and pollen as well as scarce dinocysts, ehman et al. (1976) suggested a late albian to early cenomanian age, whereas croxton (1978a, b) interpreted the sediments to be older than middle albian. in this study, the spores, pollen and brackish-water dinocysts are considered suggestive of a pre-late albian age. consequently, the kome formation is suggested to be of albian age. correlation. it has previously been stated by troelsen (1956) and rosenkrantz (1970) that the sediments exposed on the north coast of nuussuaq, east of the ikorfat fault zone, are older than the fluviodeltaic cretaceous sediments known from the rest of the nuus kome formation slibestensfjeldet formation precambrian fig. 26. contact between precambrian basement and the kome formation at talerua. the sediments onlap the relief on the basement surface. the height of the coastal cliff is c. 30 m. for location, see fig. 22. fig. 27. the type locality of the kome formation at majorallattarfik showing outcrop of facies association b (see text for explanation). person for scale (circled) is standing on delta front deposits (see also fig. 24, right-hand column). for location, see fig. 21. photo: t.c.r. pulvertaft. 45 suaq basin. however, lanstorp (1999) suggested a middle to late albian age for the atane formation at tartunaq (south-east nuussuaq), thereby shortening (or removing) the time gap between the kome and atane formations. nevertheless, the kome formation is retained as a discrete lithostratigraphic unit characterised by the onlap relationship to the basement, the coarse-grained conglomeratic character and the significant lateral facies changes (fig. 25). these features indicate a depositional environment that was markedly different from that which produced the cyclic, aggradational atane formation. in the kuuk area, the mudstones of association b and the sandstones of association c (pulvertaft 1979), referred here to the kome formation, may correlate laterally with the slibestensfjeldet and atane formations, respectively, farther to the north-west (fig. 29). thus, the mudstones of association b (fig. 24) and the mudstone-dominated unit of the type section (40–80 m in fig. 23) are thought to represent proximal correlatives of the slibestensfjeldet formation. the sandstones of association c display an erosional base and occur towards the top of the exposed sections (fig. 29). this unit contains chert pebbles, a characteristic feature of the atane formation and particularly the fluvial sandstones of the ravn kløft member of the atane formation (see below). it is tentatively suggested, therefore, that the erosional base of the sandstones in facies association c may correlate with the erosional boundary between the slibestensfjeldet and atane formations, i.e. that the uppermost kome formation in the kuuk area correlates with the lowermost atane formation to the north-west (fig. 29). cc a c a fig. 28. reference locality of the kome formation at kuuk showing outcrop of facies associations a and c (see text for explanation, see also fig. 29). for location, see fig. 21. photo: t.c.r. pulvertaft. ★ ★ ★ ★ ravn kløft mb eqalulik fmvaigat fm slibestensfjeldet fm kingittoq mb qaarsutkuuk se nw talerua ikorfat kangilia fm kome fm kome fmkome fm atane fmc a b pseudoceratium interiorense nyktericysta davisii itilli fm ikorfat fault zone ?? fig. 29. schematic profile (scale arbitrary) along the north coast of nuussuaq between kuuk and ikorfat showing correlation of the litho stratigraphic units. the sediment outcrops are bounded to the east by the basin boundary fault and to the west by the ikorfat fault zone. the occurrence of two species of dinoflagellate cysts is shown. these species are known from other localities in the nuussuaq basin. for location of outcrops, see figs 21, 22; a, b, c, facies associations discussed in the text. pale red indicates precambrian basement. slibestensfjeldet formation new formation history. the slibestensfjeldet formation has previously been described as unit 2, an informal lithological unit in the slibestensfjeldet–ikorfat area (midtgaard 1996b). name. after the prominent mountain of slibestensfjeldet on the north coast of nuussuaq (fig. 22). distribution. the slibestensfjeldet formation is known only from the north coast of nuussuaq between vesterfjeld and ikorfat (fig. 22; midtgaard 1996b). the formation is not known to be preserved north of nuussuaq, but it is expected to continue southwards, for some distance, in the subsurface. the extent of the formation west of the ikorfat fault is not known. type section. the type section is at kussinikassaq where the formation is thickest (figs 22, 30), immediately east of the basement high at ikorfat beginning at an altitude of c. 60 m a.s.l. the base of the type section is located at 70°46.20´n, 53°03.37´w. reference sections. reference sections are found at vesterfjeld, c. 860 m a.s.l., and in two gullies on slibe stensfjeldet c. 1 km east of talerua, beginning at an altitude of c. 320 m a.s.l. (fig. 30). thickness. the slibestensfjeldet formation increases in thickness from 90 m at vesterfjeld to c. 200 m at kus sinikassaq (figs 22, 30). lithology. a considerable lateral change in lithology characterises the slibestensfjeldet formation from vesterfjeld to ikorfat, a distance of approximately 10 km (fig. 30; midtgaard 1996b). at the type section, the formation is dominated by mudstones, interbedded with 0.2–2 m thick sandstone beds, which are structureless, parallel laminated, cross-laminated or hummocky cross-stratified (figs 30, 31, 32). the mudstones contain c. 5% carbon, almost entirely of organic origin, traces of pyrite are restricted to a thin coal bed, and the calculated carbon/ sulphur (c/s) ratios are c. 100. the sandstone interbeds in the c. 200 m thick mudstone unit become increasingly common upwards. at the top of this coarsening-upward succession is a medium-grained sandstone unit 5 m thick, showing steeply dipping foresets; this facies is restricted to the areas west of the talerua fault. in the east, the formation is dominated by sandstones interbedded with thin mudstones and conglomerate beds. laminated black mudstones with sand streaks are abruptly overlain by very fineto fine-grained, well-sorted sandstones dominated by horizontal lamination and hummocky crossstratification, whereas low-angle cross-bedding and wave ripple cross-lamination are subordinate (midtgaard 1996a). the sandstones are often slightly bioturbated. thin conglomerate beds composed of well-rounded clasts and showing erosional bases and wave-rippled tops, recur throughout the formation (midtgaard 1996a). fossils. the mudstones contain an assemblage of brackish-water dinocysts dominated by pseudoceratium interiorense whereas the pollen rugubivesiculites rugosus is absent (figs 29, 30). plant macrofossils from the lowermost part of the formation are referred to the ikorfat flora, which has been assigned to the early–middle albian (boyd 1998a, b, c, 2000). depositional environment. the widely distributed and thick coal bed at the base of the slibestensfjeldet formation indicates accumulation of peat under a slowly rising ground-water table. the overlying fine-grained, wave-rippled sandstones are interpreted as lake shoreline deposits and indicate an increasing rate of water level rise, which continued to the deposition of mudstones at depths below wave-base. in the mudstones, the c/s ratios of c. 100 and the stratigraphical position between the fluvial kome and the fluvial ravn kløft member indicate that the slibestensfjeldet formation was deposited in a large and deep lake. the overall, coarsening-upward, vertical facies succession is interpreted to record the infill of the lake and progradation of the shoreline. the presence of dinocysts, which are interpreted as brackish-water forms, suggests that the lake was connected to a coastal lagoon. the lake increased in depth from east to west, and wave-influenced, relatively shallow-water facies are only known from vesterfjeld and slibestensfjeldet. the conglomerate beds represent transgressive lags created by wave erosion and winnowing. the nearshore sandstones accu mulated in large complex bedforms overprinted by wave-ripples (midtgaard 1996a). the upper part of the slibe stensfjeldet formation locally developed as a facing page: fig. 30. type section of the slibestensfjeldet formation at kus sini kassaq; reference sections of the formation at vesterfjeld and slibestensfjeldet. note the marked increase in thickness towards the west (kussinikassaq). all logs are from the north coast of nuussuaq (fig. 22). for legend, see plate 1; gd, gilbert delta; ps, palaeosol. 46 vesterfjeld kussinikassaqslibestensfjeldet vaigat fm gd c. 50 m not shown ps gd 20 m 10 0 ps mud sand pebbles sl ib es te ns fje ld et f or m at io n sl ib es te ns fje ld et f or m at io n a ta ne f or m at io n sl ib es te ns fje ld et f or m at io n r av n k lø ft m em be r r av n k lø ft m em be r r av n k lø ft m em be r mud sand pebbles mud sand pebbleskome fm kome fm 47 48 fig. 31. mudstones with thin sandstone beds in the lower part of the slibestensfjeldet formation at kussinikassaq, near ikorfat. for location, see fig. 22. fig. 32. the upper part of the slibestensfjeldet formation in its type locality at kussinikassaq, where heterolithic sandstone bodies with wavegenerated sedimentary structures (see also midtgaard 1996 a, b) are overlain by lower shoreface sandstones (see fig. 30) and locally by gilbert delta foreset beds. the slibestensfjeldet formation is erosionally overlain (dashed line) by the fluvial ravn kløft member of the atane formation. the height of the section is c. 100 m. for location, see fig. 22. slibestensfjeldet formation ravn kløft member atane formation gilbert delta in the ikorfat area (figs 30, 32; midtgaard 1996b). the extent of the lake or lagoon west of ikorfat is not known. it may be speculated that subsidence along the ikorfat fault affected the development of the lake. boundaries. the lower boundary of the slibestensfjeldet formation is sharp and varies very little laterally; it is placed at the base of a 20−110 cm thick, continuous coal bed that overlies a palaeosol (fig. 30). the upper boundary of the slibestensfjeldet formation with the atane formation is defined by an abrupt contact between shoreface sandstones or gilbert delta sandstones and coarse-grained, pebbly fluvial sandstones of the ravn kløft member (figs 30, 32). the boundary has been interpreted either as a minor angular unconformity (gry 1940; koch 1964) or as the conformable, yet erosional base of a channel (ehman et al. 1976; schiener 1977; croxton 1978a, b; midtgaard 1996b). midtgaard (1996b) interpreted the boundary as a sequence boundary. geological age. based on plant macrofossils referred to the ikorfat flora, the slibestensfjeldet formation has been assigned an early to middle albian age (boyd 1998a). the dinocysts suggest a late albian age (ehman et al. 1976; croxton 1978a, b) or more likely a pre-late albian age (this study). correlation. the slibestensfjeldet formation in the slibestensfjeldet outcrops may correlate with coarsening-upward successions in the kome formation at kuuk (fig. 29). outcrops connecting these two areas are, however, not present. upernivik næs formation revised formation history. the sedimentary outcrops at upernivik næs were described by steenstrup (1883b) who noted that the coalbearing sediments have a thickness of at least 860 m. the lithology was not considered different from the coalbearing successions known from nuussuaq and disko. the macrofossil plants collected at upernivik næs by steenstrup were referred to the atane flora by heer (1883b). in contrast, seward (1926) compared them to the kome flora. koch (1964) considered the fossil flora from upernivik ø as difficult to interpret due to similarities to both the kome and the atane floras and suggested that the upernivik ø flora could occupy an intermediate position. troelsen (1956) did not recognise the upernivik næs formation, which was later described by koch (1964). name. the formation is named after upernivik næs on the south-western corner of the island of upernivik ø (figs 2, 33). distribution. the formation is known from upernivik ø. the marginal marine sediments on the island of qeqertarsuaq and on itsaku are here referred to the upernivik næs formation (figs 2, 73; ødum & koch 1955, christiansen et al. 2000; j.g. larsen & pulvertaft 2000). type section. the type section is located on the south-western corner of upernivik ø, along a large stream that flows out on the south coast of the island c. 2.5 km east of upernivik næs (figs 33–35). the base of the type section is located at 71°09.88´n, 52°54.52´w. reference section. a reference section is exposed in a coastal cliff on western qeqertarsuaq (figs 2, 36, 73). thickness. the formation is c. 1600 m thick at upernivik næs, where the strata generally dip c. 15° towards the north-east (henderson & pulvertaft 1987). croxton (1978b) measured a 1310 m long section (termed the c9 section) along the southern coast of upernivik næs; this section coincides with the type section in its upper part. in the type section, the formation is exposed up to c. 500 m a.s.l. (fig. 33). lithology. the section at upernivik næs is dominated by sandstones; mudstones constitute only a minor part (figs 34, 35). on qeqertarsuaq, sandstone-dominated successions are locally overlain by clast-supported conglomerates, in which the clasts are dominantly from the metamorphic rocks of the widely exposed karrat group (fig. 37; henderson & pulvertaft 1987). the formation has been studied at a reconnaissance level by midtgaard (1996b), who divided the sediments of the formation into four facies associations. facies association a. coarse-grained, poorly sorted sandstones constitute up to 34 m thick composite depositional units (fig. 38). some of these are distinctly fining upwards. basal channel lags contain pebbles of crystalline rocks and locally also mudstone clasts, logs and carbonaceous debris. the sandstones are cross-bedded with most palaeocurrent directions towards the western quadrant. soft-sediment deformation structures are common. 49 50 facies association b. well-sorted, fineto mediumgrained sandstones with abundant mudstone clasts characterise the association. siltstone and mudstone beds and laminae form an important component and in places are heterolithic and ripple-cross laminated. the sandstones are dominated by planar or trough cross-bedding that indicates palaeocurrents to the north. some sandstones have mud-draped foresets and show bundled lamination (fig. 39). the mud-dominated heteroliths and mudstones contain much comminuted plant debris; bioturbation varies from moderate to intense (fig. 39). facies association c. sandstones interbedded with mudstones constitute 3−8 m thick successions. the sandstones range from very fineto very coarse-grained, the beds are erosionally based and pinch out laterally. they are cross-laminated, cross-bedded or locally horizontally stratified. the tops of the sandstone beds are moderately bioturbated. the mudstones are black, fissile and sandstreaked, and sometimes contain abundant plant debris and vitrinite lenses. sandstone dykes are common, and gentle folding indicates that they were intruded before or during compaction. facies association d. coarsening-upward successions comprising black fissile mudstones, mudstones with sandstone lenses, mud-draped ripple cross-laminated, fine-grained sandstones, and trough cross-bedded sandstones with comminuted plant debris (midtgaard 1996b). these deposits constitute only a minor part of the section shown in fig. 34. fossils. the upernivik næs formation contains angio sperm plants referred to the upernivik næs flora (or upernivik flora), which is interpreted as intermediate in age between the kome flora and the atane flora (koch 1964; boyd 1998a). macrofossil plants from qeqertarsuaq correspond to species known from outcrops at upernivik næs and atanikerluk (ødum & koch 1955). few palynomorphs have been recorded from croxton’s (1978b) c9 section through coastal exposures and the type section at upernivik næs. the presence of a few specimens of the pollen rugubivesiculites rugosus throughout the section indicate a middle albian to turonian age, and the co-occurrence with a few specimens of the dinocysts nyktericysta davisii and quan touendinium dictyophorum in the lower part of the section suggests correlation to the lower part of the atane formation (?ravn kløft member). a large specimen of the dinocyst nyktericysta davisii has been recorded from the middle part of the c9 section. similar forms are common in sediments on qeqertarsuaq (christiansen et al. 2000) and from a single fully marine sample from the atane formation at ikorfat dated as early turonian. the sediments in facies 71°09' 71°12' 53° 2 km upernivik næs glacier 740 1760 1939 10 10 15 16 22 20 18 3020 20–25 sea/lake upernivik næs formation with conglomerate dolerite/basalt fault (ticks on downthrow side)■■ ■■ ■■ ■ ■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ fig. 33. geological map of upernivik næs on the south-western tip of upernivik ø. the type section of the upernivik næs formation is located in the ravine 2.5 km east of upernivik næs. slightly modified from henderson & pulvertaft (1987). for location, see fig. 2. the contour interval is 200 m. facing page: fig. 34. type section of the upernivik næs formation; from midtgaard (1996b). notice the presence of tidal deposits. for legend, see plate 1; for location, see fig. 33. 0 clay gravelsi sand clay gravelsi sand clay gravelsi sand clay gravelsi sand 50 100 150 500 550 m 350 400 450 200 ?? 250 300 51 association b contain trace fossils such as skolithos, are nicolites, conichnus, planolites, rhizocorallium and pelecypodichnus (midtgaard 1996b). depositional environment. the sandstones of facies association a are interpreted to have been deposited in braided fluvial channels. the scarcity of interbedded floodplain or levee facies suggests that the channels were migrating laterally. the abundant soft-sediment deformation structures suggest synsedimentary tectonic activity. the heterolithic, burrowed sandstones (facies association b) are interpreted to have been deposited in an estuarine environment – in tidal channels, tidal point bars, tidal deltas, tidal flats, bay mudstones, and on the shoreface as transgressive lags. the palaeocurrents are interpreted to reflect longshore tidal transport. the interbedded sandstones and mudstones (facies association c) are thought to represent a coastal plain environment and include lagoonal, mudflat and marsh deposits, crevasse splay deposits and deposits in small channels. the scarcity of rootlets and the presence of trace fossils suggest brackish-water conditions. intrusion of the sandstone dykes may have been triggered by earthquakes. the coarsening-upward successions (facies association d) are interpreted as deltaic deposits. the lack of wave-generated sedimentary structures in facies association d in the upernivik ø section indicates deposition in a low-energy environment, probably an interdistributary bay or the inner part of an estuary, i.e. a bayhead delta (midtgaard 1996b). higher-energy facies on qe qertarsuaq and itsaku suggest the influence of waves and storms on the delta deposits (j.g. larsen & pulvertaft 2000). boundaries. the lower boundary of the upernivik næs formation is not exposed in the type section nor in the reference section. conglomerates with large clasts of both weathered and fresh basement rocks have been reported close to the boundary fault at upernivik næs (fig. 33), from a small outcrop on the north-western part of upernivik ø (henderson & pulvertaft 1987 fig. 46), from qeqertarsuaq and from itsaku (rosenkrantz & upernivik næs formation precambrian karrat group fig. 35. type locality of the upernivik næs formation. the central peak is c. 1700 m high. for location, see fig. 33. 52 53 pulvertaft 1969). the boundary between the kome and the upernivik næs formations is at present not known. the upper boundary of the formation is an erosional contact with the kangilia formation on itsaku and with quaternary deposits on upernivik næs and qeqertarsuaq. geological age. the upernivik næs flora contains angio sperm plants, which suggests a greater similarity to the atane flora than to the kome flora (koch 1964). koch recommended, however, that until the cretaceous floras are thoroughly revised, the upernivik næs flora should fig. 37. clast-supported conglomerate of the upernivik næs formation on qeqertarsuaq. hammer for scale; for location, see fig. 73. fig. 36. outcrop of the upernivik næs formation on qeqertarsuaq; for location, see fig. 73. reddish-brown intrusions cross-cut (right) and cap the sediments. 54 be eliminated from the discussion of the age of the kome and atane formations and treated as an independent problem (koch 1964 p. 540). rosenkrantz (1970) suggested that the upernivik næs flora is of albian−turonian age. boyd (1998a) suggested that the upernivik næs flora is contemporaneous with his ravn kløft flora which he referred to the middle–late albian and earliest cenomanian. the poor palynological assemblage indicates a late albian to early turonian age (croxton 1978a, b; this study). fig. 38. channellised sandstones of the upernivik næs formation in the lower part of the type section. thickness of sandstone unit is approximately 10 m. for location, see fig. 33. fig. 39. detail of tidally influenced deposits of the upernivik næs formation in the lower part of the type section. person for scale; for location, see fig. 33. 55 correlation. the upernivik næs flora was suggested to be coeval with the ravn kløft flora by boyd (1998a). this suggests a correlation between the upernivik næs formation and the ravn kløft member of the atane formation (figs 13, 16). atane formation redefined formation history. nordenskiöld (1871) erected the atane formation (atanelagren), naming it after the now-abandoned settlement of atane (now ataa, figs 3, 40). he described the unit to be more shaly than both the underlying kome formation and the overlying upper atanikerdluk for mation, and interpreted the atane formation as freshwater deposits. he referred the formation to the upper cretaceous, following heer (1868), with a distribution on southern nuussuaq (between atanikerluk and ataa), northern disko (at qullissat) and at kuuk on northern nuussuaq (at altitudes above 250–300 m). outcrops with coal seams were visited by steenstrup (1874), who noted the occurrence of marine invertebrates in mud stones between the coal seams at paatuut and nuuk qiterleq. the atane formation normally appears as a striped white, grey and black sandy and shaly unit rich in coalified plant fossils (fig. 41) that formed the basis of the classic studies by heer (1868, 1870, 1874a, b, 1880, 1883a, b). however, self-combustion of the mudstones has locally produced brick-red, hard and fissile burnt slabs (fig. 42) in which fossils are excellently preserved as impressions. burnt mudstones are fairly common in the paatuut (formerly spelled patoot, pâtût or pautût) and kingittoq areas, and they were included in the regional collection that steenstrup sent to o. heer (steenstrup 1883a, b). based on their colour, heer treated the samples of burnt mudstones from paatuut as lithostratigraphically different from those containing the atane flora. thereby a patoot formation (i.e. the sediments comprising the patoot flora) was introduced and subsequently treated as a formal unit (troelsen 1956). steenstrup (1883c) objected to heer’s interpretation of a stratigraphic boundary between the unburnt and the burnt mudstones from paatuut since he had observed that the burnt mudstones are laterally equivalent to the normal, unburnt sediments. steenstrup (1883a) suggested that all the cretaceous sediments from svartenhuk halvø to disko should be enclosed in one lithostratigraphic unit. he envisaged the depositional environment of the horizontally bedded mudstones and sandstones as shallow marine on account of his collections of invertebrates (bivalves and echinoids), although the well-preserved plant remains indicated deposition close to vegetated areas (steenstrup 1883a p. 48). the early cretaceous to miocene age proposed by heer for the coal-bearing deposits in west greenland was discussed by a. heim and j.p.j. ravn, who both suggested that heer’s interpretation spans too long a period. instead they interpreted all the sediments as upper cretaceous to eocene (heim 1910), more specifically the kome flora as albian, the atane flora as cenomanian, the patoot flora as senonian and the atanikerdluk flora as eocene (ravn 1918 p. 320). ravn (1918) envisaged the plant bearing sediments as having been deposited in a freshto brackish-water environment, in which the rate of deposition equalled the rate of subsidence. the inoceramids provided evidence of marine conditions and ravn therefore concluded that at times subsidence exceeded deposition and marine environments were established. furthermore, he observed no changes in the style of sedimentation during the cretaceous and inferred that deposition had occurred continuously, presumably within a restricted period and consequently at relatively high rates of sediment accumulation. troelsen (1956) maintained the distinction between the kome, atane and patoot formations, quoting descriptions by nordenskiöld (1871) and heer (1883b). these authors suggested that the kome and atane formations are separated by a slight angular unconformity (troelsen 1956). koch (1964) regarded the pautût formation (sensu heer 1883b) as an artificial unit. henderson et al. (1976 fig. 303) quoted the lithostratigraphy of troelsen (1956), but stressed that the kome, atane and pautût formations lacked formal stratigraphic definition and referred all the fluvial and deltaic cretaceous sediments on disko and southern to central nuussuaq to the atane formation. ehman et al. (1976) used the term atane formation informally for the cretaceous non-marine deposits and found that the sediments range in age from albian through santonian. g.k. pedersen & pulvertaft (1992) referred all nonmarine cretaceous sediments on disko and nuussuaq to the atane formation. they claimed (p. 263) that the term atane formation “had outlived the other formation names that have been used in the past for the different isolated outcrops of cretaceous non-marine strata in west greenland”. based on the data available at that time, g.k. pedersen & pulvertaft (1992 p. 263) interpreted these strata as “belonging to the same deposi56 atane formation maligât formation vaigat formation atane formation volcanic scree q 12 5 m fig. 41. typical outcrop of the atane formation at paatuut on southern nuussuaq. note the prominent cyclic alternation of lithologies. for location, see fig. 40. q, quikavsak formation; compare with fig. 14. ataa ivissussat ivissussat qaqqaat 247701 247901 247801 1010 paatuut uppall giesec monum 53 °w 45 ’ 21 ’ 21 ’ 18 ’ 70 °1 5’n 15 ’ 18 ’ 12 ’ 30 ’ 45 ’ 53 ° paatuutkløften ippigaarsukkløften a t a a t a k u u a ivisaannguit fig. 40. geological map of the south coast of nuussuaq from saqqaqdalen to ivisaannguit, simplified from a.k. pedersen et al. (2007b). the type localities of the atane, quikavsak and atanikerluk formations are all found within this area. the wells ggu 247701, ggu 247801 and ggu 247901 indicated on the map were drilled to obtain technical information on the coal-bearing strata of the atane formation. contour interval is 200 m; for location, see fig. 2. tional system and the same age interval, so that no more than one formation name seems at present to be required for the entire non-marine cretaceous of the area”. later, from sedimentological studies, midtgaard (1996b) argued for the continued distinction between the atane, kome and upernivik næs formations, as well as for the new slibestensfjeldet formation. boyd (1993) considered, mostly on floral evidence, that the sediments at paatuut should not be included in the atane formation. he therefore retained the pautût flora and the pautût formation (boyd 1993 p. 253). there are, however, no lithological or sedimentological arguments for retaining the pautût formation of troelsen (1956) or boyd (1993), and we therefore recommend that it be abandoned and included in the atane formation. name. from a former settlement ataa (old spellings atane, atâ) on the south coast of nuussuaq immediately west of ataata kuua (figs 2, 3, 40; nordenskiöld 1871 plate xxi). today only ruins of a few turf houses are seen at ataa. distribution. the atane formation is restricted to disko and nuussuaq. it comprises the lowermost exposed deposits on eastern disko and southern and central nuussuaq. marginally marine successions of albi an– cenomanian age in the northern parts of the nuussuaq basin are referred to the upernivik næs formation. exposures of the atane formation are found east of the disko gneiss ridge and east of the kuugan nguaq–qunnilik fault (fig. 10; chalmers et al. 1999). on northern nuussuaq, the atane formation is present east of the ikorfat fault zone where it overlies the slibestensfjeldet formation (fig. 22; midtgaard 1996b). the distribution of the atane formation west of the outcrop area is little known. there are, however, no sedimentological data from the atane formation to indicate that the disko gneiss ridge formed a morphological element during deposition of the formation. finds of sandstone xenoliths with chert pebbles in lava flows on western disko have been suggested to be derived from the atane formation (chalmers et al. 1999). the xenoliths could, however, also have been derived from the itilli formation, which occurs on western nuussuaq. type section. the type section of the formation is the cored well ggu 247801 (70°19.87´n, 52°55.18´w) in the gorge of the ataata kuua river (figs 14, 40) where these sediments were first described by nordenskiöld (1871). the next geologist on the site was k.j.v. steen strup. he noted in 1872 that the description given by nordenskiöld did not match the outcrop and that the 57 umiusat nunngarut qallorsuaq keglen 5 km kingittoq s a q q a q d a l e n 12 ’ 09 ’ 09 ’ 06 ’ 03 ’ 03 ’ 06 ’ 15 ’ 15 ’ 52 °3 0’w intrusion atanikerluk formation quikavsak formation kangilia formation atane formation kingittoq member atane formation qilakitsoq member atane formation undivided undifferentiated pre-quaternary deposits precambrian ice sea/lake boreholevaigat formation maligât formation quaternary cover abandoned settlement iviangernat vaigat 70 °n 70 ° tartunaq takisut qeqertaq aasivik aqqullugiaq quikassaap kuua naajaat atanikerluk qallunnguaq geographical position was inaccurate (steenstrup 1874). he suggested that nordenskiöld had measured the type section of the atane formation south-east of the ataa river, at “the northern end of the patoot gorges” (steen strup 1883b p. 64). later rosenkrantz (1970 p. 422) stated that: “... the type locality atâ, 14 km south-east of the old village atâ on the south coast of nûgssuaq, is rather close to pautût..”, but in henderson et al. (1976) the type locality of the atane formation was again given at atâ (ataata kuua). the exposures along the western and eastern slopes of the ataata kuua gorge constitute a long, well-exposed section through the atane formation, and this has been measured photogrammetrically (fig. 15; a.k. pedersen et al. 2007b). the sedimentological log representing the type section (fig. 43) was measured in the continuously cored borehole ggu 247801, which reached a depth of 566 m below terrain (76 m below sea level). reference sections. skansen, j.p.j. ravn kløft, kingittoq and qilakitsoq provide well-exposed reference sections for the atane formation (figs 48, 51, 53, 60). the two continuously cored boreholes ggu 247701 and ggu 247901, both from the paatuut area, are reference boreholes for the atane formation. they comprise strata from the qilakitsoq member and have been described in detail by ambirk (2000). thickness. the atane formation occurs in a number of fault blocks (j.m. hansen 1980b; chalmers et al. 1999; marcussen et al. 2001, 2002). the faults are, however, rarely exposed and although they have been recognised on seismic profiles acquired in vaigat and uummannaq fjord (marcussen et al. 2002), it has not been possible to determine their throw. furthermore the lower boundary of the formation is very rarely exposed. therefore the thickness of the atane formation can only be given as an estimate. the thickest exposed sections occur at ivissussat on the south coast of nuussuaq (700 m) and at qilakitsoq on central nuussuaq (800 m). seismic data from the south coast of nuussuaq show an unconformity at a depth of 1.5−2 km that has been suggested to represent the base of the atane formation (christiansen et al. 1995; chalmers et al. 1999). a seismic line across vaigat indicates a minimum thickness of 3000 m for the atane formation (fig. 44; marcussen et al. 2001, 2002). lithology.the atane formation comprises mudstones, heteroliths, sandstones and coal beds; the relative proportion of the lithologies varies regionally. the formation is characterised by aggradation of 10–40 m thick depositional cycles, which often coarsen upwards (fig. 45). the sandstones are typically fineto medium-grained, well sorted or heterolithic, with various types of crossfacing page: fig. 43. type section of the atane formation (and reference section for the qilakitsoq member) from ataata kuua, borehole ggu 247801. this core documents the stratigraphic interval known from the well-exposed section in the ataata kuua river gorge and extends this section down to 76 m below sea level. tr, transgressive sandstones. for legend see plate 1; for location, see fig. 40. burnt shales fig. 42. well-exposed outcrop at paatuut showing the self-combusted burnt shales (red) that grade into the typical black-and-white unburnt sediments of the atane formation (qilakitsoq member). height of section is approximately 50 m. for location, see fig. 40. 58 59 ggu 247801 550 500 450 400 350 300 250 200 150 100 50 0 m clay silt sand pebbles vf f m c vc clay silt sand pebbles vf f m c vc clay silt sand pebbles vf f m c vc tr tr tr tr tr tr 247815 ★ ★ ★ ★ ★ 60 bedding, numerous soft-sediment deformation structures and local bioturbation. the sandstones consist of quartz, variably kaolinised feldspars, fragments of older quartz-rich sedimentary rocks, and small amounts of mica and other detrital minerals together with comminuted plant debris. trace fossils are locally abundant. the sandstones are cemented by carbonate or clay minerals but are generally friable. pervasive cementation is only patchily developed. the sandstones have sheet or ribbon geometries. the mudstones are weakly laminated, typically silty, with kaolinite as the dominant clay mineral accompanied by a little mica. the mudstones fall into two groups: (1) mudstones ranging from silt-streaked mudstone to 0 750sp 500 250 0 ns 1 km c. 0.5 km quaternary sediments cretaceous sediments 500 1000 t w t ( m se c) 1500 fig. 44. seismic section across the vaigat (for location, see fig. 2; sea level is at twt=0). the submarine section continues into exposures of the formation up to 200–400 m a.s.l. on both sides of vaigat indicating a minimum thickness of the atane formation of 3000 m (from marcussen et al. 2002). sp, seismic shot points. df dp dp ch ch df d ts fig. 45. outcrop of the atane formation in the paatuut area showing the four depositional environments characterising the atane formation: delta front (df), delta plain (dp), fluvial or distributary channel (ch) and transgressive shoreface deposits (ts). the delta front deposits form coarsening-upward units, indicated by triangular symbols. the dashed lines trace the bases of delta front mudstones. the sedimentary succession is cut by a dyke (d). the thickness of the section is c. 120 m; for location, see fig. 40. wave-rippled sand streaked mudstone to heterolithic sandstone, and (2) mudstones with plant debris, ranging from carbonaceous mudstone to clayey coal. the coal beds are interbedded with carbonaceous mudstones, sand-streaked mudstones and heterolithic cross-laminated sandstones with local root horizons. the coal beds are typically less than 0.8 m thick. in several horizons, coal balls are observed, and silicified wood fragments are common. the coals belong to the ‘banded coal’ type dominated by the lithotypes vitrain and clarain. macerals of the vitrinite group are often well preserved owing to the prevailing low rank of the coal. detailed study of the organic particles and their geochemistry permit a distinction between freshand brackish-water environments of coal deposition (shekhar et al. 1982; bojesen-koefoed et al. 2001; g.k. pedersen et al. 2006). the atane formation is characterised by depositional cycles that are typically 10–40 m thick (figs 41, 46). the simplest cycles are seen on eastern disko where they consist of 10–30 m thick sandstone sheets separated by c. 5 m thick units of mudstone and coal beds. on southern and central nuussuaq, the cycles range in thickness from 5–15 m up to 40 m. they include coarseningupward units beginning with mudstones at the base passing up into heterolithic mudstones, then heterolithic sandstones with hummocky cross-stratification and ripple cross-lamination and finally into cross-bedded sandstones, often with coal debris. the typical coarseningupward cycle is capped by carbonaceous mudstones with thin sandstone or coal beds (pedersen & pulvertaft 1992 fig. 5). bioturbated sandstones form part of many cycles. in the paatuut area, 32 such cycles have been identified (olsen 1991; dueholm & olsen 1993). fossils. the fossils of the atane formation comprise macroflora, spores and pollen, dinocysts, and invertebrate fossils (heer 1868, 1870, 1874a, b, 1880, 1883a, b; ravn 1918; koch 1964; birkelund 1965; ehman et al. 1976; croxton 1978a, b; boyd 1990, 1992, 1993, 1994, 1998a, b; olsen & pedersen 1991; koppelhus & pedersen 1993; mcintyre 1994a, b, c; nøhr-hansen 1996; lanstorp 1999; dam et al. 2000). macrofossil plants from the atane formation are referred to three floras. the ravn kløft flora is characterised by angiosperms and conifers, and has few species in common with the atane flora (boyd 1998a, b). the atane flora is a mixture of older cretaceous species together with well-differentiated angiosperm species representative of many families (koch 1964; k.r. pedersen 1976). the paatuut flora is dominated by conifer and angiosperm leaf species, and has many species in common with the atane flora as well as numerous endemic species (koch 1964; boyd 1992, 1993, 1994). spores and pollen occur in most mudstone samples from the atane formation, although often the preservation is poor or the number of identifiable specimens is low. regional studies of the spore and pollen assemblages have been carried out by ehman et al. (1976) and croxton (1976, 1978a, b). the latter distinguished two stratigraphically important assemblages, the first without angiospermous (tricolpate) grains (supposed to be cenomanian or older), the second with angiospermous grains but lacking both aquilapollenites and complex triporate grains. the second assemblage was proposed to be older than late campanian (croxton 1978a p. 76). a large number of samples from the cretaceous were examined on a reconnaissance basis by mcintyre (1994a, b, c, personal communication 1997) whereas koppelhus & pedersen (1993) and lanstorp (1999) studied fewer sections in greater detail. marine dinocysts were described from central nuussuaq (ilugissoq, qilakitsoq, tunoqqu, southern part of agat dalen) and from southern nuussuaq (paatuut, ataata kuua, nuuk qiterleq and nuuk killeq) by nøhr-hansen (1996), croxton (1978a, b), mcintyre (1994a, b, c, personal communication 1997) and dam et al. (2000). brackish-water dinocysts of late albian to cenomanian age have been described from northern nuussuaq (figs 29, 30; nøhr-hansen in sønderholm et al. (2003), asuk and in the f93-3-1 core from kuugannguaq, northern disko. a sparse fauna of marine invertebrates (echinoids and bivalves) is known from southern and central nuussuaq including sphenoceramus patootensis, sphenoceramus pinniformis, and oxytoma tenuicostata (ravn 1918; olsen & pedersen 1991). ammonites from a single horizon at alianaatsunnguaq were referred to the scaphites ventricosus – inoceramus involutus zone by birkelund (1965); the atane formation in agatdalen has yielded baculites codyensis (reeside) and radially ribbed inoceramids of the steenstrupi species-group, an assemblage referred to the clioscaphites montanensis zone (birkelund 1965; dam et al. 2000). ammonites from the atane formation occur as redeposited fossils in the itilli formation on central nuussuaq (dam et al. 2000). clioscaphites saxitonianus septentrionalis (birkelund) has been described from ilugissoq, and at tunoqqo clioscaphites sp. aff. saxitonianus (mclearn) occurs together with a single scaphites cf. svartenhukensis (birkelund 1965; dam et al. 2000). although koch (1964) suggested that the marine fossils are restricted to a few horizons, olsen & pedersen (1991) reported that the marine fossils recur through 61 62 tr tr tr tr tr tr tr tr tr tr tr tr tr tr tr tr tr c. 2 km c. 1 km 50 m fig. 46. simplified sedimentological logs from the ivissussat–paatuut area; colours depict depositional environments, see plate 1. the logs document the cyclic depositional pattern and local variations in thickness and lateral changes of facies. individual phases of delta progradation may be correlated over short distances between neighbouring sections. the resulting high-resolution lithostratigraphy has only local significance. for legend, see plate 1; tr, transgressive sand. for location, see fig. 40. 63 the formation in the marine mudstones at the base of the delta front cycles. boyd (1993) also reported marine invertebrates from several stratigraphic horizons. trace fossils are locally abundant and include dac tyoloidites ottoi, ophiomorpha nodosa, o. irregulaire, taenidium serpentinum, planolites isp., teichichnus isp., thalassinoides isp., diplocraterion isp., skolithos isp. (fürsich & bromley 1985; g.k. pedersen & rasmussen 1989; bojesen-koefoed et al. 2001; g.k. pedersen & bromley 2006, bromley & g.k. pedersen 2008). depositional environment. the atane formation is interpreted as having been deposited in a major delta system, and four depositional environments are distinguished: the prograding marine delta front, fluvial or distributary channels of the delta plain, lakes and swamps of the delta plain, and the marine shoreface. the coarsening-upward successions of mudstones, heteroliths and fine-grained sandstones are interpreted as delta front deposits. in some successions, they include interdistributary bay deposits (j.m.hansen 1976; midt gaard & olsen, 1989; midtgaard 1991; olsen 1991; olsen & pedersen 1991; g.k. pedersen & pulvertaft 1992; dueholm & olsen 1993, nielsen 2003) (figs 45, 46). they are interpreted to have resulted from progradation of shelf deltas sensu elliott (1989). olsen (1993) described channel mouth complexes from the uppermost part of many delta front cycles at paatuut. the fluvial sandstone sheets on eastern disko are interpreted as sandy braided river deposits (johannessen & nielsen 1982; koppelhus & pedersen 1993; bruun 2006). these may constitute the multi-storey fill of fluvial valleys on south-east nuussuaq (jensen 2000; jensen & pedersen in press). fluvial sandstones with ribbon geometry are interpreted as slightly sinuous distributary channel deposits (olsen 1991, 1993). the carbonaceous mudstones, sandstones and coal seams are interpreted as freshwater lake or swamp deposits on the delta plain (midtgaard 1991; olsen & pedersen 1991; g.k. pedersen & pulvertaft 1992; koppelhus & pedersen 1993; nielsen 2003; møller 2006; g.k. pedersen et al. 2006). this facies association is interpreted as representing the vertical aggradation of a subaerial to shallow limnic, upper and lower delta plain. the thin beds of sand with marine trace fossils are interpreted as transgressive sand sheets, the erosional base of which constitutes a ravinement surface (midtgaard 1991; olsen & pedersen 1991; g.k. pedersen & rasmussen 1989). olsen (1991) and dueholm & olsen (1993) documented 32 cycles of delta progradation in the paatuut area and comparable numbers of delta cycles are known from the qilakitsoq area. j.m. hansen (1976) observed that the deltaic cycles in a vertical section at any locality show a remarkable similarity, indicating an aggradational stacking pattern. correlation of individual deltaic cycles is only possible between closely spaced outcrops (fig. 46), but lateral changes in depositional environment within the atane formation may be demonstrated on a larger scale (fig. 16). boundaries. on disko, the atane formation overlies a basement high in the fp-93-3-1 well (fig. 19). on nuussuaq, the base of the atane formation is exposed along the north coast where the formation erosively overlies the slibestensfjeldet formation between ikorfat and vesterfjeld (figs 30, 32). the boundary has been interpreted either as a minor angular unconformity (gry 1940; koch 1964) or as a conformable, erosional base of a channel (ehman et al. 1976; schiener 1977; croxton 1978a, b; midtgaard 1996b). midtgaard (1996b) interpreted the unconformity as a sequence boundary (fig. 32). the upper boundary of the atane formation is a marked erosional unconformity (see fig. 16). on eastern disko and nuussuaq, it is overlain by fluvial and lacustrine paleocene deposits of the atanikerluk for mation. on northern disko, the atane formation is erosively truncated by the itilli formation at kussinerujuk and asuk and by the volcanic vaigat formation at naajannguit (hald & pedersen 1975; a.k. pedersen 1985). along the south coast of nuussuaq, the atane formation is unconformably overlain by the marine kangilia formation (dam et al. 2000), by incised valley fills of the quikavsak formation (dam et al. 1998a; dam & sønderholm 1998; dam et al. 2000; dam 2002), and by the eqalulik formation. in the aaffarsuaq valley the atane formation is unconformably overlain by the itilli formation (dam et al. 2000) with rare ammonites (birkelund 1965). west of ilugissoq (central nuussuaq), the upper boundary of the atane formation is not exposed. on northern nuussuaq, the upper boundary of the formation can only be seen between vesterfjeld and ikorfat but is generally poorly exposed. directly east of the ikorfat fault zone, the atane formation is overlain by mudstones with late campanian to maastrichtian ammonites (birkelund 1965), referred to the itilli for mation. between vesterfjeld and ikorfat, the atane for mation is overlain by the eqalulik formation or by volcanic rocks of the vaigat formation (figs. 22, 29; a.k. pedersen et al. 1996, 2006b). 64 geological age. the scarcity of marine fossils combined with the long range of many terrestrial plants, spores and pollen makes it difficult to determine the age of the atane formation with precision; for this purpose not only the presence but also the abundance of palynomorphs is important. the age of the formation is bracketed by the palynomorphs or plant macrofossils in the underlying kome and slibestensfjeldet formations and by dinocysts and ammonites in the overlying itilli formation on central and northern nuussuaq (fig. 16). the oldest parts of the atane formation are albi an– cenomanian and early turonian and occur on disko, on southern nuussuaq at kingittoq and eastwards, and on northern nuussuaq around j.p.j. ravn kløft (croxton 1978a, b; koppelhus & pedersen 1993; mcintyre 1994a, b, c, personal communication 1997; boyd 1998a, b; lanstorp 1999; this study). sediments of albian–ceno manian to turonian age have been reported from asuk, and at kussinerujuk the atane formation is overlain by the cenomanian kussinerujuk member of the itilli formation (mcintyre 1994a, b; bojesen-koefoed et al. 2007). the presence of coniacian deposits is based on the occurrence of scaphites ventricosus at aliana atsun nguaq (birkelund 1965). the youngest parts of the formation are of late santonian to earliest campanian age and occur on southern and central nuussuaq (paatuut, qilakitsoq, agatdalen; olsen & pedersen 1991; boyd 1992, 1993, 1994; nøhr-hansen 1996; d.j. mcintyre, personal communication 1997; dam et al. 2000). these sediments are dated from ammonites (clioscaphites montanensis zone), dinocysts and macroplant fossils (birke lund 1965; boyd 1992, 1993, 1994; nøhr-hansen 1996; dam et al. 2000). coniacian to early santonian am mo nites in the itilli formation on central nuussuaq overlying the atane formation indicate that the latter is santonian or older in this area (birkelund 1965; dam et al. 2000). correlation. the atane formation is in part coeval with the marginally marine upernivik næs formation north of nuussuaq, and with the lower part of the deep marine itilli formation on nuussuaq (west of the kuugan nguaq−qunnilik fault) and on svartenhuk halvø (figs 13, 16). subdivision. the atane formation is divided into four members: the albian–cenomanian skansen member and ravn kløft member, the albian to lower turonian kingittoq member, and the upper turonian to santonian – lowermost campanian qilakitsoq member, which includes the ?lower campanian itivnera bed. skansen member new member history. the skansen member includes the sediments referred to the atane formation on southern and eastern disko. the coal beds at skansen were visited by giesecke in 1807 and 1811 (in: steenstrup 1910). he described the alternation between sandstones and coal beds, and noted that spherical concretions (kieskugeln) are frequent in the sandstones. the history of coal mining at skansen was summarised by k.j.v. steenstrup, who also measured a geological profile of the coastal cliff at skansen (steenstrup 1874 plate viii). name. the name is taken from the former settlement skansen (old spelling skandsen) or aamaruutissat, on the south coast of disko (fig. 124). the name ‘skansen’ (meaning rampart or palisade) originates from a thick and very prominent sill of columnar jointed basalt, named innaarsuit in greenlandic. distribution. the skansen member crops out on southern and eastern disko (figs 124, 132; a.k. pedersen et al. 2000, 2001, 2003). the sediments are exposed in stream sections along the coast and in the inland valleys such as kvandalen and laksedalen. the skansen member is not known from wells and its distribution outside the area of outcrop is unknown. the cretaceous sediments recorded in seismic sections in disko bugt south and east of disko (chalmers et al. 1999) presumably in part belong to the skansen member. type section. the type section of the skansen member is on the south-facing slope behind the settlement of skansen (figs 47, 48, 124). the sedimentology and palynology of this section have been studied in some detail (koppelhus & pedersen 1993; bruun 2006; møller 2006a, b). the base of the type section is located at 69°26.40´n, 52°26.32´w. reference sections. reference sections are found at illunnguaq (koppelhus & pedersen 1993) and pingu (fig. 132). thickness. only the upper 470 m of the skansen member are exposed. however, geophysical data indicate that sediments with a thickness of c. 2 km are present east of disko (chalmers et al. 1999), suggesting that the skansen member may reach a considerable thickness in the subsurface. 65 lithology. the skansen member is dominated by white to yellow sandstones with sheet geometry, intercalated with coal seams, mudstones and heteroliths rich in plant debris arranged in a cyclic depositional pattern (figs 47, 49). sixteen cycles are recognised, typically 20–30 m thick (bruun 2006). the sandstones consist of quartz, more or less kaolinised feldspars, fragments of older quartz-rich sedimentary rocks, and small amounts of mica and other detrital minerals together with comminuted plant debris. the sandstones are generally friable, since carbonate or clay mineral cement is only locally developed. the sandstones are mediumto coarsegrained, in places with pebbly channel lags of intraformational mudstones, coal clasts or pebbles of chert and ordovician limestone (a.k. pedersen & peel 1985). they are dominantly cross bedded or structureless (fig. 49), but commonly show soft-sediment deformation structures. sandstone sheets with a width: thickness ratio in excess of 15 are characteristic. the interbedded finegrained sediments are dark grey to black due to the ubiquitous presence of comminuted plant debris. facies vary rapidly both vertically and laterally between heterolithic, cross laminated sand, mudstone with plant debris, massive brownish mudstone and coal beds, most of which are less than 1 m thick. locally, root horizons or rare tree stumps are present. the coal beds are c. 0.5 m thick, and the coals are subbituminous, with up to 20% siliciclastic particles and very little pyrite. the coals are dominated by huminite which originates from wood and plant tissue rich in cellulose (møller 2006a, b). fossils. macroflora, spores and pollen are known from the skansen member (heer 1883a, b; seward 1926; miner 1932a, b, 1935; ehman et al. 1976; croxton 1978a, b; koppelhus & pedersen 1993). the plant fossils from six localities (innanguit, killusat, skansen, pingu, ujara sussuk, illukunnguaq) were referred to the atane flora by heer (1883b p. 93). assemblages of palynomorphs d d fig. 47. type locality of the skansen member of the atane formation at skansen, southern disko (for location, see fig. 124). the skansen member is dominated by fluvial sandstones, interbedded with fine-grained floodplain deposits and thin coal seams. the peak is at 470 m a.s.l. and the sediments are cut by several dykes (d). 66 have been studied at reconnaissance level at marraat (southern disko) and skansen (ehman et al. 1976), at kuuk quamasoq and pingu (croxton 1978a, b; d.j. mcintyre, personal communication 2003) and in more detail at skansen and illunnguaq (koppelhus & pedersen 1993) (fig. 124). depositional environment. the sandstone sheets in the skansen member are interpreted as sandy braided river deposits. the sediment transport directions on eastern disko vary from north-north-west at pingu to southwest at gule ryg and skansen (g.k. pedersen & pulvertaft 1992 fig. 6), suggesting the presence of a huge alluvial cone with its apex east of disko. apparently, the disko gneiss ridge did not affect the depositional pattern of the skansen member (johannessen & nielsen 1982; g.k. pedersen & jeppesen 1988; koppelhus & pedersen 1993; bruun 2006). the carbonaceous mudstones, sandstones and coal seams are interpreted as freshwater lake or swamp deposits representing the vertical aggradation of a subaerial to shallow, limnic floodplain to upper delta plain. the mud sand mud sand 180 200 250 mm 50 0 110 130 fig. 48. type section of the skansen member; note the dominance of fluvial channel deposits, (see also figs 47 and 49). base of section is at c. 120 m a.s.l. for legend, see plate 1. 67 spores and pollen represent vegetation dominated by conifers and ferns; there are no indications – neither palynological evidence nor the presence of pyrite – to suggest marine or brackish-water conditions (g.k. pedersen & pulvertaft 1992; koppelhus & pedersen 1993; møller 2006). boundaries. the lower boundary of the skansen member is not exposed. the upper boundary is an erosional unconformity overlain by the paleocene atanikerluk formation. at pingu, the skansen member is overlain by fluvial sand of the akunneq member (figs 16, 132), while it is overlain by lacustrine mudstones of the assoq member at tuapaat on southern disko. the presence of late cenomanian or turonian strata at the top of the illunnguaq section was suggested by koppelhus & pedersen (1993), but was not confirmed by examination of additional samples. the upper boundary of the skansen member corresponds to the upper boundary of the atane formation in the area where the skansen member occurs. geological age. the skansen member is dated on the basis of spores and pollen. at its type locality, a midcenomanian age seems most likely, with a maximum age range of late albian to cenomanian for the palynomorph assemblages from this member (croxton 1978a, b; ehman et al. 1976; koppelhus & pedersen 1993). correlation. the fluvial skansen member is laterally equivalent to the deltaic kingittoq member on northern disko (between qullissat and kussinerujuk), and on southern nuussuaq (between atanikerluk and kingittoq) (fig. 16). the skansen member also correlates with the fluviodeltaic ravn kløft member on northeastern nuussuaq and in part at least with the upernivik næs formation on upernivik ø. palynomorphs are not preserved in the lower part of the itilli formation on svartenhuk halvø and on western nuussuaq, which precludes firm correlation between the skansen member and the itilli formation (figs 13, 16). ravn kløft member new member history. the ravn kløft member is part of the cretaceous succession overlying the slibestensfjeldet formation on north-eastern nuussuaq. these sediments have previously been referred to the atane formation (nordenskiöld 1871; gry 1942; midtgaard 1996b) or to the upernivik næs formation (rosenkrantz 1970; henderson et al. 1976). name. the member is named after the gorge j.p.j. ravn kløft where it forms impressive outcrops (fig. 50a, b). j.p.j. ravn studied the geology of the nuussuaq basin in 1909, with focus on the marine invertebrates from the region (ravn 1918). distribution. the ravn kløft member is known from outcrops along the north coast of nuussuaq between ikorfat c. 20 m fp c fig. 49. cross-bedded mediumto coarsegrained fluvial sandstones of the skansen member at skansen. note the intraformational conglomerate (c) composed of mudstone and coal clasts (lower part of photo), and the thin interval of carbonaceous mudstone (floodplain deposits; fp) in the upper part. 68 c. 6 0 m kingittoq member ravn kløft member fig. 50. a: type locality of the ravn kløft member on the west-facing slopes of the j.p.j. ravn kløft gorge on slibestensfjeldet. note the very thick fluvial sandstone bodies at the top of the ravn kløft member that can be traced westwards to ikorfat. the ravn kløft member is overlain by the kingittoq member (atane formation). for location, see fig. 22. b: large-scale cross-bedding in sandstones of the thick, amalgamated fluvial channel deposits at the top of the ravn kløft member at ravn kløft (c. 395 m in fig. 51a). the height of the exposure is c. 3 m. a b 69 and qaarsut (fig. 22). its wider distribution is not known due to lack of exposures. type section. the eastern slope of j.p.j. ravn kløft constitutes the type section (figs 22, 51a). the base of the type section is located at 70°45.18´n, 52°55.08´w. reference sections. reference sections are found in stream exposures just east of ikorfat, at slibestensfjeldet and at vesterfjeld (figs 22, 51b, 52). thickness. the thickness of the ravn kløft member is about 450 m in the type section, thinning both towards the east and the west (midtgaard 1996b). the basal pebbly sandstone varies in thickness from less than 2 m to a maximum of 56 m. the overlying deposits are c. 400 m thick in j.p.j. ravn kløft. lithology. a variety of lithologies and sedimentary facies are present in the ravn kløft member. the member is tripartite, comprising a lower pebbly sandstone unit, a middle unit of mudstones, heteroliths and fine-grained sandstones with coarsening-upward or fining-upward depositional patterns, and an upper unit of thick-bedded mediumto coarse-grained sandstones interbedded with mudstones and heteroliths (fig. 50b; midtgaard 1996b). the lower unit (c. 35–62 m on fig. 51a) includes a c. 3 m thick conglomeratic sandstone with rounded pebbles and large clasts of coaly mudstone, fossil wood and mudstone, interpreted as a channel lag. this is overlain by pebbly coarse-grained sandstones with large-scale cross-bedding, followed by coal-bearing mudstones with beds consisting of small scaidopityoides leaves (midtgaard 1996b). the cross-bedding indicates unidirectional northward sediment transport. the middle unit (c. 62–340 m on fig. 51a) comprises a variety of facies. coarsening-upward successions comprise laminated mudstones, heterolithic sandstones and hummocky cross-stratified sandstones, which locally are capped by trough cross-bedded sandstones, mudstones with root horizons and thin coal beds. wave-ripple crests are oriented ese–wnw. comminuted car bonaceous debris is abundant. fining-upward successions comprise mediumto coarse-grained, well-sorted sandstones with cross-bedding, double mud-drapes, and abundant reactivation surfaces. they are overlain by heteroliths with current and wave ripple cross-lamination followed by black, laminated mudstones. synaeresis cracks are common. the upper unit (c. 340–400 m on fig. 51a) of the ravn kløft member consists of thick, erosively based bodies of greyish sandstone alternating with dark heterolithic mudstone. the sandstones are fineto coarsegrained or conglomeratic, and dominantly cross-bedded. foresets are often oversteepened and large-scale soft-sediment deformation structures are very common (fig. 52). composite sandstone bodies may be up to 60 m thick (fig. 50), and may extend laterally for at least 8 km. the cross-bedding indicates sediment transport to the ne–n–nw. the sandstone bodies alternate with thinly interbedded rippled sandstones, laminated mudstones and heteroliths containing abundant rootlets and plant fossils, thin coal beds and palaeosols at certain levels. fossils. a brackish-water dinocyst assemblage dominated by nyktericysta davisii has been identified in a number of the mudstone beds. the pollen rugubivesiculites rugosus has its first stratigraphical occurrence within the member (fig. 29). depositional environment. the lower unit of the ravn kløft member is interpreted as a basal fluvial valley-fill conglomerate overlain by fluvial sandstones deposited by unidirectional northward-flowing currents. the middle units are interpreted as interbedded deltaic and tidal estuarine deposits, and wave-ripple crestlines suggest an ese–wnw-oriented coastline. the overlying single to multi-storey sandstone bodies represent a variety of fluvial styles from slightly sinuous single channels to braided rivers with multiple channels and northward palaeocurrents. the presence of rootlets, vitrinite lenses and local coal beds in the floodplain sandstones and mudstones indicates the existence of a range of sub-environments from subaerial floodplain to shallow-water swamps or ponds adjacent to the fluvial channels. the fluvial sandstones show an increasing tendency up-section to amalgamate and form thick, multi-storey sandstone sheets (midtgaard 1996b). boundaries. the lower boundary of the ravn kløft member is the same as for the atane formation. the thickness variation of the lower fluvial sandstones of the ravn kløft member suggests that the base had a relief of nearly 55 m on a regional scale (midtgaard 1996b). the upper boundary is placed at the top of a c. 60 m thick, amalgamated multi-storey fluvial sandstone sheet which is abruptly overlain by a succession of interbedded mudstones, heteroliths, sandstones and thin coal seams referred to the kingittoq member (fig. 50). geological age. based on the palynomorphs, the ravn kløft member is assigned a late albian – early ceno 70 sl ib es te ns fje ld et f or m at io n a ta ne f or m at io n sl ib es te ns fje ld et f m r av n k lø ft m em be r a ta ne f or m at io n r av n k lø ft m em be r r av n k lø ft m em be r r av n k lø ft m em be r k in gi tt oq m em be r k in gi tt oq m em be r 0 50 100 mud sand mud sand mud sand mud sand mud sand 150 200 250 m 0 50 450 m 300 350 400 100 r av n k lø ft m em be r 150 200 250 a b fig. 51. a: type section of the ravn kløft member at j.p.j. ravn kløft. b: reference section at ikorfat. the thick sandstone beds at the top of the member are overlain by the kingittoq member. for legend, see plate 1; for location, see fig. 22. 71 manian age (ehman et al. 1976; croxton 1978a, b; this study). a distinct macroflora (the ravn kløft flora) of middle albian to early cenomanian age was established by boyd (1998a, b, c, 2000); this flora was broadly contemporaneous with the upernivik næs flora of rosen krantz (1970). correlation. the ravn kløft member is laterally equivalent to the skansen member in southern and eastern disko and the older parts of the kingittoq member on central and southern nuussuaq and northern disko. the ravn kløft member is possibly coeval with the lower part of the itilli formation on western nuussuaq and svartenhuk halvø and with parts of the upernivik næs formation on upernivik ø (fig. 16). kingittoq member new member history. the kingittoq member includes the sediments referred to the atane formation on northern disko, south-eastern nuussuaq, as well as some of those on northern nuussuaq. 0 mud base 495 m a.sl. sand gravel 20 30 10 0 mud sand gravel 20 30 m m 10 40 m se nw tidal–estuarine lacustrine fluvial valley fill fig. 52. two closely spaced sections through the ravn kløft member showing the rapid lateral facies changes especially of the fluvial channel deposits. the sections correspond approximately to the interval 255–285 m in fig. 51a. for legend, see plate 1; for location, see fig. 22. 72 name. the member is named from the coastal slopes at kingittoq on the south coast of nuussuaq, where it is well exposed (fig. 40). distribution. the kingittoq member is known from outcrops on southern nuussuaq from saqqaqdalen to kingittoq, on northern disko from qullissat to kussinerujuk, and on northern nuussuaq from vesterfjeld to ikorfat (figs 2, 22, 40, 50a). 0 clay silt sand pebbles vf f m c vc f m c clay silt sand pebbles vf f m c vc f m c clay silt sand pebbles vf f m c vc f m c 140 130 120 160 170 180 190 200 210 220 230 240 250 260 270 280 290 300 310 320 330 340 350 360 370 380 390 400 410 420 430 440 450 460 470 480 490 m 150 110 100 90 80 70 60 50 30 40 20 10 fig. 53. type section of the kingittoq member of the atane formation at kingittoq, south coast of nuussuaq (for location, see fig. 40). note the higher proportion of delta channel deposits compared to the qilakitsoq member (fig. 60). compare with figs 54 and 55 to appreciate the facies variations within the member. marine dinocysts are only found in a few of the delta front units. for legend, see plate 1. 73 type section. the type section of the kingittoq member is exposed in the slopes at kingittoq on southern nuussuaq (figs 40, 53, 59). the base of the type section is located at 70°09.77´n, 52°31.38´w. reference sections. reference sections are found in the upper reaches of j.p.j. ravn kløft and at asuk (figs 54, 55). thickness. the thickness of the kingittoq member is not known, but it probably exceeds 1 km. in the area of the type section, more than 600 m are exposed (a.k. pedersen et al. 2007a), c. 450 m are exposed at kuussinerujuk and up to 150 m are exposed above the ravn kløft member on northern nuussuaq. lithology. on southern nuussuaq, the kingittoq member is characterised by 10−25 m thick depositional cycles, which comprise coarsening-upward successions of mudstones, heteroliths and well-sorted sandstones (figs 53–58). complex sandstone sheets, up to 40 m thick, are often overlain by carbonaceous heterolithic mudstones with thin coal beds and constitute fining-upward successions (figs 53, 59). in the atanikerluk area, the complex sandstone sheets constitute c. 40% of the section (nielsen 2003) whereas their proportion is higher in saqqaqdalen (shekhar et al. 1982) and at kussinerujuk (pulvertaft & chalmers 1990). the member is characterised by aggradational stacking of the depositional cycles (figs 57, 58). the sandstones are pale, friable, mediumto coarsegrained, in places with pebbly channel lags of intraformational mudstone or coal clasts, and are either crossbedded or structureless, commonly with soft-sediment deformation structures (midtgaard 1991; olsen 1991; g.k. pedersen & pulvertaft 1992; jensen 2000; nielsen 2003; g.k. pedersen et al. 2006). the mudstones are grey, silty, carbonaceous, at several localities with high c/s ratios, and are weakly laminated (nielsen 2003). the mineralogy is similar to that in the rest of the atane formation. the coal beds are interbedded with mudstones with plant debris, sand streaked mudstones and cross laminated sand. root horizons are seen locally, but preserved tree stumps are rare. a few of the coal beds have been studied in detail (shekhar et al. 1982; bojesenkoefoed et al. 2001; g.k. pedersen et al. 2006). coars ening-upward successions of grey mudstones, wave-rippled heterolithic sandstones and fine-grained sandstones with swaley and hummocky cross stratification are especially well developed at asuk (fig. 56). the sedimentary structures are enhanced by drapes of comminuted plant (coal) 0 40 m 30 20 10 mud sand gravel fig. 54. reference section of the kingitoq member at j.p.j. ravn kløft (for location, see fig. 22). note the predominance of delta plain facies deposited in shallow lakes, small channels and interdistributary bays. see also fig. 51a. base of section at c. 650 m a.s.l. for legend, see plate 1. 74 debris. marine fossils are found locally and bioturbation is generally slight. thin beds of well-sorted, structureless sandstone are seen at kingittoq, whereas similar beds at asuk show erosional bases and contain numerous trace fossils, of which ophiomorpha nodosa is prominent (fig. 55). fossils. plant macrofossils, spores and pollen are known from the kingittoq member (heer 1883a, b; seward 1926; ehman et al. 1976; croxton 1978a, b; lanstorp 1999). heer (1883b) described plant fossils collected from four outcrops of the kingittoq member (qullissat, asuk, qallunnguaq, and atanikerluk) and referred these to the atane flora and the patoot flora. marine dinocysts are known from the kingittoq member (d.j. mcintyre, personal communication 1997) and finds of rare individuals of three marine invertebrates, nucula cancellata, nucula sp., and lucina occidentalis, were reported from kingittoq (ravn 1918). the trace fossils ophiomorpha nodosa, taenidium serpentinum, teichichnus rectus, thalassinoides isp. occur locally in the kingittoq member. 10 0 20 30 40 50 60 70 80 90 100 110 120 130 140 m clay silt sand vf f m c vc clay silt sand vf f m c vc k in gi tt oq m em be r a ta ne f or m at io n k us si ne ru ju k m em be r it ill i f or m at io n 0 10 20 m tr tr tr tr tr tr tr fig. 55. reference section of the kingittoq member at asuk, north coast of disko (for location, see fig. 2). the section is domi nated by stacked delta front deposits and intervals assigned to the delta plain facies association are thin. note that only one thin coal seam is present (44 m). see also figs 56, 57 and 80. at this locality, the member is clearly more influenced by marine deposition than at kingittoq (fig. 53). tr, transgressive sandstones. for legend, see plate 1. 75 depositional environment. the kingittoq member is interpreted to represent fluvial to deltaic deposits and is referred to four facies associations: the channel and delta plain associations (dominant), the delta front or mouth bar association (subordinate) and the transgressive sand sheet association (rare) (midtgaard 1991; g.k. pedersen & pulvertaft 1992; dam et al. 2000; jensen 2000; nielsen 2003; g.k. pedersen et al. 2006). the cross-bedded, mediumto coarse-grained sandstones are interpreted as fluvial channel deposits. the sedimentary structures frequently indicate downstream accretion on bars and rare development of point bars, indicating that the fluvial channels were mostly braided. the complex sheets of coarse-grained sandstones are interpreted as having been deposited as the multi-storey fill of fluvial valleys (jensen 2000; jensen & pedersen in press). the carbonaceous mudstones, sandstones and coal beds are interpreted to represent the vertical aggradation of subaerial to freshwater lacustrine or swamp deposits. the high c/s ratios indicate that deposition occurred in freshwater environments on the upper delta plain (g.k. pedersen & pulvertaft 1992; nielsen 2003). the coal beds are interpreted as the in situ accumulations of plant remains in freshwater environments (bojesen-koefoed et al. 2001; g.k. pedersen et al. 2006). this facies association is subordinate at asuk. the coarsening-upward heteroliths are interpreted as delta front or mouth bar deposits formed during progradation of shelf deltas (g.k. pedersen & pulvertaft 1992; bojesen-koefoed et al. 2001; nielsen 2003; g.k. pedersen et al. 2006). this facies association is dominant at asuk. the thin beds of sand with marine trace fossils are interpreted as transgressive sand sheets (midtgaard 1991). the erosive base constitutes a ravinement surface. in the kingittoq member, the transgressive sand sheets are best developed at asuk. the kingittoq member is dominated by stacked upper delta plain deposits in the atanikerluk area and at qullissat, and by lower delta plain and stacked delta front deposits at asuk. boundaries. the lower boundary of the kingittoq member is exposed on northern nuussuaq where the kingittoq member overlies the ravn kløft member (fig. 50). here the boundary is a drowning surface which separates the fluvial sandstones from the overlying delta plain deposits. in the rest of the nuussuaq basin, the lower boundary is not exposed. the upper boundary, corresponding generally to the upper boundary of the atane formation in the outcrop area of the member, is everywhere an erosional unconformity overlain by the itilli, quikavsak, atanikerluk and vaigat formations (fig. 16). on the north coast of delta front s m fig. 56. delta front succession of the kingittoq member at asuk (27–35 m in fig. 55). note the upward increase in frequency and thickness of sandstone beds. the well-sorted sandstones with hummocky and swaley cross-stratification (s) are indicative of a high-energy depositional environment. the mudstone (m) at the top of the section is interpreted as an interdistributary bay deposit. for location, see fig. 2. 76 nuussuaq, the upper boundary of the kingittoq member is generally poorly exposed due to landslides and rock glaciers, but east of ikorfat it is overlain by small outcrops of campanian–maastrichtian mudstone (the itilli for mation) and paleocene tuffaceous mudstone (the eqalulik formation) (birkelund 1965; a.k. pedersen et al. 2006b) (fig. 22). ts delta front channel channel 25 m delta front delta front delta front ts ts ts fig. 57. the kingittoq member at asuk showing thick fluvial channel deposits overlying coarsening-upward delta front successions capped by transgressive shoreface sandstones (ts). for location, see fig. 2 and fig. 55: 44–83 m. channel channel channel delta plain delta plain delta plain delta plain delta front delta front delta front channel 30 m fig 58. outcrop of the kingittoq member on the eastern slope of the qallunnguaq valley (for location, see fig. 40). delta plain facies constitute a large proportion of the succession and delta front deposits are thin and fine-grained suggesting lowenergy depositional environments (lower delta plain with swamps, shallow lakes and interdistributary bays (nielsen 2003)). the channel sandstone at the top of the photo graph is the lowermost fluvial unit in a multi-storey sandstone body studied by jensen (2000) and jensen & pedersen (in press). 77 geological age. the geological age of the kingittoq member ranges from albian (tartunaq) to cenomanian (atani kerluk, saqqaqdalen, kingittoq, kussinerujuk, asuk and ravn kløft) and into early turonian (saqqaqdalen, kingittoq and asuk). the deposits are dated on the basis of spores, pollen and marine dinocysts (croxton 1978a, b; lanstorp 1999; mcintyre 1994a, personal communication 1997; this study). correlation. the kingittoq member is coeval with the skansen member (eastern disko); the kingittoq member of southern nuussuaq may be equivalent the ravn kløft member on northern nuussuaq (fig. 16). the kingittoq member correlates with the upernivik næs formation on upernivik ø, and possibly to the lower part of the itilli formation of western nuussuaq and svartenhuk halvø. qilakitsoq member new member history. the qilakitsoq member includes sediments referred to the atane formation on southern and central nuussuaq, and sediments formerly referred to the now abandoned pautût (or patoot) formation (see above under history of the atane formation). name. the name is derived from the qilakitsoq stream, a tributary to the kuussuaq river which flows through the aaffarsuaq valley on central nuussuaq (fig. 82). distribution. the qilakitsoq member is the sole representative of the atane formation in central nuussuaq and parts of southern nuussuaq. outcrops are found from agatdalen to ilugissoq along the north slope of the aaffarsuaq valley and along the south coast from paatuut to alianaatsunnguaq (figs 2, 40, 82). type section. the type section of the qilakitsoq member is along the qilakitsoq stream on central nuussuaq (figs 60, 82). the base of the type section is located at 70°27.97´n, 53°27.13´w. reference sections. reference sections are found at nuuk killeq and ataata kuua (type section of the atane formation; fig. 43). two continuously cored boreholes from the paatuut area (ggu 247701 and ggu 247901) provide reference sections of the qilakitsoq member and have been described in detail by ambirk (2000). thickness. the type section documents a minimum thickness of 480 m, but correlation with nearby sections shows that at least 820 m are exposed in the qilakitsoq area, and c. 600 m are known from the well at ataata kuua together with nearby outcrops (fig. 43; a.k. pedersen et al. 2007b). lithology. the qilakitsoq member is characteristically cyclic, individual cycles typically passing up from mudstones through heteroliths, well-sorted sandstones, coarser grained sandstones with ribbon geometry, and finally into carbonaceous mudstones with coal beds overlain by thin sheets of bioturbated sandstone (fig. 60; midtgaard sill df ch ch atane formation atanikerluk formation fig. 59. the kingittoq member at king ittoq, see 310–465 m in fig. 53. most fluvial channels are braided, but a point bar succession is seen in the centre of the photo with inclined accretionary surfaces (indicated by dotted line); ch, fluvial channel; df, delta front. for location, see fig 40. 78 1991; olsen 1991, 1993; g.k. pedersen & pulvertaft 1992; boyd 1993; dueholm & olsen 1993; ambirk 2000; dam et al. 2000). the coarsening-upward successions are typically 5−25 m thick but may reach up to c. 70 m (fig. 61). they can be traced laterally through closely spaced outcrops for up to 8 km in the paatuut area (fig. 46; olsen 1991, 1993; dueholm & olsen1993) and also in the qilakitsoq area. the mudstones at the base of the coarsening-upward units are dark grey, silty and weakly laminated, and may contain marine fossils (olsen & pedersen 1991; nøhrhansen 1996). these grade up into sand streaked mudstones and heterolithic sandstones, where wave-ripples and swaleyand hummocky cross-stratification are enhanced by drapes of comminuted plant debris (fig. 60, 33–50 m). the successions are frequently topped by medium-grained, trough cross-bedded sandstones (fig. 60, 145-158 m). bioturbation may locally be very intense and totally obscure primary structures. 0 10 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 0 190 200 210 220 230 180 clay silt sand pebbles vf f m c vc f m c clay silt sand pebbles vf f m c vc f m c clay silt sand pebbles vf f m c vc f m c 240 250 260 270 280 290 300 310 320 350 340 330 360 370 380 390 400 410 420 430 450 460 470 480 m 440 fig. 60. type section of the qilakitsoq member (atane formation) on the western side of the qilakitsoq stream, central nuussuaq. note the higher proportion of delta front deposits compared to the kingittoq member (fig. 53). for location, see fig. 82; for legend, see plate 1. 79 other sandstones are pale, friable, mediumto coarsegrained, in places with channel lags of pebbles or intraformational clasts of mudstone and coal. these sandstones are cross bedded or structureless, commonly with softsediment deformation structures. they form single-storey sandstone ribbons with a width:thickness ratio as low as 6 and a sinuosity of 1.2 (olsen 1993) (fig. 60: 0–11 m). successions of carbonaceous heteroliths interbedded with cross laminated sand, sand streaked mudstones, mudstones with plant debris and coal seams occur in all outcrops of the qilakitsoq member. their dark grey to black colour reflects the abundance of comminuted plant debris. locally, root horizons or rare tree stumps are present. numerous coal beds at paatuut and ataa are more than 0.8 m thick. concretions and ‘coal balls’ are observed in several horizons, as well as silicified wood. dirt bands in coal beds are common, mainly consisting of carbonaceous mudstones (fig. 60: 20–32 m). sheets of structureless or strongly bioturbated sandstone with an erosional base occur in most outcrops of the qilakitsoq member. they range from thin (c. 0.05 m) fine-grained sandstones to thick (c. 3 m) mediumto coarse-grained sandstones. upwards, these sandstones may pass gradually into mudstones (fig. 60, 180–182 m). fossils. plant macrofossils, spores and pollen are known from the qilakitsoq member (heer 1883a, b; seward 1926; ehman et al. 1976; croxton 1978a, b; boyd 1990, 1992, 1993, 1994). heer (1883b) described plant fossils collected from three outcrops of the qilakitsoq member (alianaatsunnguaq, ataa, and paatuut) and referred these to the atane and patoot floras. boyd (1992, 1993) recognised three or more floral communities (nearshore lacustrine, backswamp, levee and riparian vegetation) in the fossil flora from the paatuut area (the paatuut flora). marine dinocysts have been identified in samples from the qilakitsoq member (d.j. mcintyre, personal communications 1997, 1999; nøhr-hansen 1996). the sparse marine invertebrates include ammonites (cliosca phites saxitonianus septentrionalis (birkelund), clioscaphites sp. aff. saxitonianus (mclearn), scaphites ventricosus, scaphites cf. svartenhukensis, and baculites codyensis (reeside)) (birkelund 1965; dam et al. 2000); bivalves (sphenoceramus patootensis, s. pinniformis, oxytoma tenuicostata) and echinoderms (ravn 1918; olsen & pedersen 1991; boyd 1993). trace fossils are common in the qilakitsoq and nuuk killeq areas. burrows such as planolites isp., teichichnus isp., dactyolidites ottoi, and helminthopsis horizontalis occur in the coarsening-upward successions, whereas ophiomorpha nodosa and o. irregulaire are seen in the sandstone sheets (fig. 62; g.k. pedersen & rasmussen 1989; dam et al. 2000; g.k. pedersen & bromley 2006). the trace fossil assemblages indicate the presence of both suspension and deposit feeders. depositional environment. the qilakitsoq member is interpreted as being constructed of aggradational deltaic deposits referred to four facies associations: the marine delta front association, the distributary channel association, the delta plain association and the transgressive sand sheet association (midtgaard 1991; olsen 1991, 1993; g.k. pedersen & pulvertaft 1992; dueholm & olsen 1993; dam et al. 2000). atane formation 65 m vaigat formation sh ch delta front sh l fig. 61. outcrop of the qilakitsoq member on the western slope of gully immediately west of qilakitsoq. the member is dominated by stacked delta front successions. the outcrop corresponds to the interval 200–380 m in fig. 60; ch, channel; l, lagoon; sh, shoreface. note that the uppermost delta front succession is abnormally thick at this locality. for location, see fig. 82. 80 the marine delta front association comprises coarsening-upward units where mudstones with marine fossils deposited below storm-wave base are overlain by heterolithic mudstones and sandstones which pass up into shallower water facies. these units are interpreted as the result of delta front progradation of shelf deltas. the distributary channel association consists of crossbedded, mediumto coarse-grained sandstones, which are interpreted as having been deposited within almost straight to slightly sinuous fluvial channels (olsen 1991, 1993). small fluvial channels are included in the delta plain association, which is characterised by carbonaceous mudstones, sandstones and coal beds, interpreted as having formed through vertical aggradation of subaerial to freshwater lacustrine or swamp deposits on the lower and upper delta plain. the coal beds are interpreted as the in situ accumulations of plant remains in both freshand brackish-water environments. the transgressive sand sheet association includes thin beds of sand with erosive bases and often with numerous marine trace fossils, indicating deposition on a marine shoreface (g.k. pedersen & rasmussen 1989; midtgaard 1991; olsen 1991; olsen & pedersen 1991; g.k. pedersen & pulvertaft 1992; g.k. pedersen & bromley 2006). the erosional bases represent ravinement surfaces. boundaries. the lower boundary of the qilakitsoq member is not exposed. the upper boundary, which corresponds to the upper boundary of the atane for mation, is everywhere an erosional unconformity. dif ferent units overlie the member throughout the region; these include the itilli, kangilia, quikavsak and atanikerluk formations as well as volcanic rocks of the vaigat formation (fig. 16). geological age. the geological age of the qilakitsoq member is middle turonian to santonian at ataata kuua, coniacian at alianaatsunnguaq (birkelund 1965), san tonian − earliest campanian at paatuut, nuuk killeq and on central nuussuaq (koch 1964; boyd 1990, 1992; olsen & pedersen 1991; g.k. pedersen & pulvertaft 1992; nøhr-hansen 1996; d.j. mcintyre, personal communication 1997). stable carbon isotopes in wood fragments suggest a middle to late santonian age (ambirk 2000). correlation. the qilakitsoq member correlates with part of the itilli formation (western nuussuaq and svartenhuk halvø). subdivision. the qilakitsoq member includes the itivnera bed in the aaffarsuaq valley near tunoqqu. itivnera bed revised bed history. the itivnera beds were described by dam et al. (2000) as fluvial sandstones incised into the top of the qilakitsoq member (atane formation) at two widely separated localities: itivnera and ataata kuua (figs 40, 82). the erosional, lower boundary was interpreted as a sequence boundary. at itivnera, the fluvial deposits are fig. 62. the trace fossil ophiomorpha irregulare is characteristic of transgressive shoreface sandstones in the qilakitsoq member, especially at qilakitsoq and nuuk killeq. for location, see fig. 2. 81 overlain by deep marine deposits belonging to the aaffarsuaq member of the itilli formation. at ataata kuua, a valley incised into the atane formation is filled by turbidite mudstones, sandstones and conglomerates referred to the late maastrichtian to earliest paleocene kangilia formation; the itilli formation is not present at ataata kuua (figs 14, 15, 16). the qilakitsoq member here is of turonian–santonian age. fine-grained organicrich sediments preserved in the top of a small fluvial channel contain palynomorphs indicating a coniacian age. the fluvial channel deposits at ataata kuua that were previously referred to as itivnera beds are now assigned to the qilakitsoq member on the basis of photogrammetric mapping (a.k. pedersen et al. 2007b) and dating of the fine-grained channel fill, which lies within the age range of the qilakitsoq member at this locality. the presence of early campanian fluvial deposits at ataata kuua cannot be demonstrated. the itivnera beds sensu dam et al. (2000) constituted the basal part of the itilli formation. in the present paper, this unit is restricted to the cemented channel sandstones at itivnera which are defined here as the itivnera bed of the qilakitsoq member. name. the strata are named after the saddle feature named itivnera between nalluarissat and tunoqqu (fig. 82). the name was derived from a nearby locality shown on the geodetic map from 1966 (geodætisk institut 1966); the spelling has not been changed to modern greenlandic orthography because the locality is not shown on later geodetic maps. fig. 63. type locality of the itivnera bed in central nuussuaq (for location, see fig. 82). the bed comprises three channellised sandstone units up to 38 m thick that cut down into santonian deltaic deposits of the qilakitsoq member (atane formation). the itivnera bed is overlain by submarine fan deposits of the itilli formation (aaffarsuaq member). the sandstone cliffs are about 16 m high and the spacing between the valleys is less than 100 m (see dam et al. 2000). cl si sand pb 0 5 10 15 m fig. 64. type section of the itivnera bed. neither the base nor the top of the fluvial sandstone unit is exposed (from dam et al. 2000). for legend, see plate 1; for location, see fig. 82. 82 type section. the strata exposed on the south-facing slope of nalluarissat between aaffarsuaq and kangersooq, just west of itivnera, are designated as the type section (figs 63, 64). the type section is located at 70°29.50´n, 53°08.87´w. distribution. the bed has only been recognised on the south-facing slope of nalluarissat between aaffarsuaq and kangersooq (fig. 82). thickness. the bed is up to 38 m thick and confined to lensoid bodies up to 100 m wide. lithology. the bed consists of cross-bedded coarse-grained sandstones, arranged in fining-upward successions form lensoid bodies arranged like pearls on a string (figs 63, 64). basement pebble conglomerates are locally present. the sediments between the cemented sandstone bodies are covered by scree (dam et al. 2000). fossils. macrofossils have not been found. depositional environment. the sandstones of the itivnera bed were deposited in fluvial channels (dam et al. 2000). boundaries. the itivnera bed erosively overlies deltaic deposits of the qilakitsoq member (atane formation). in the type section, the fluvial sandstones are succeeded by turbiditic mudstones and sandstones of the aaffarsuaq member (itilli formation). the lower boundary is no longer interpreted as a sequence boundary because there is insufficient evidence that a sea-level fall preceded the rise in sea level that marks the transition to the deep-water deposits of the aaffarsuaq member. geological age. at nalluarissat, just west of itivnera, the qilakitsoq member is late santonian in age, and at tunoqqu, immediately east of itivnera, the aaffarsuaq member is of early to middle campanian age (nøhrhansen 1996). deposition of the fluvial sandstone bodies at itivnera is therefore well constrained to the early campanian. itilli formation new formation history. the strata exposed in river sections in the itilli valley on western nuussuaq were informally assigned to the itilli formation by j.m. hansen (1980b). the itilli formation is here extended to include the unnamed upper turonian to campanian marine strata on svartenhuk halvø and nuussuaq (cf. birkelund 1965; henderson et al. 1976; j.g. larsen & pulvertaft 2000). furthermore, on northern disko at kussinerujuk and asuk, outcrops previously correlated with the paleocene kangilia formation (j.m. hansen 1980b; pulvertaft & chalmers 1990) are here assigned to the itilli formation (see below). the itivnera beds of the itilli formation (dam et al. 2000) are, however, now redefined as the itivnera bed of the qilakitsoq member (atane for mation). on itsaku (svartenhuk halvø), the ?upper campa nian/maastrichtian to paleocene succession has been suggested to be equivalent either to both the itilli and kangilia formations (i.e. campanian to paleocene) or to the kangilia formation alone (i.e. upper maastrichtian to paleocene), based on correlation of two major conglomerate horizons with tectonic events recognised on nuussuaq (j.g. larsen & pulvertaft 2000). based on zircon provenance data, this succession is here assigned to the kangilia formation (see below). name. after itilli, a major valley transecting nuussuaq from north-west of marraat on the south coast to west of niaqornat on the north coast (figs 2, 65). distribution. the itilli formation is exposed in the itilli valley (fig. 65) and on the north coast of nuussuaq in the ravines between ikorfat and niaqornat (fig. 74) where it has been drilled in the shallow wells ggu 400705, ggu 400706, and ggu 400407 (chri stiansen et al. 1994a), and the formation is probably also present in the fp94-11-02, fp94-11-04 and fp94-11-05 wells (dam & nøhr-hansen 1995). it is also well exposed on central nuussuaq along the slopes of the valley of aaffarsuaq between qilakitsoq and tunoqqu, along the slopes of the valley kangersooq (fig. 82), and in the valley of agatdalen including the shallow well ggu 400702 (fig. 113; nøhr-hansen 1996; dam et al. 2000). on disko, the formation is exposed at asuk and kussinerujuk. facing page: fig. 65. map of the southern part of the itilli valley showing outcrops of the itilli formation (anariartorfik member) and the eqalulik formation and location of the wells marraat-1, ganw#1, gane#1, gank#1 gro#3 and fp94-9-01. based on rosenkrantz et al. (1974) and hald (1976). contour interval 200 m. for location, see fig. 2. 83 qaasersut 782 nuusaq v a i g a t kuussuaq ganw#1 fp94-9-01 marraat-1 gro#3 gane#1 gank#1 niaqornaarsuk marraat killit eqalulik gassøen 54° 70°30’ 888 1084 1088 795 830 1136 u k a l e r s a l i k sermersalik i l u g i s s o r s u a q a n a r i a t o r f i k 5 km contour interval 200 m i t i l l i p i n g u n n g u u p k u u a well maligât formation vaigat formation eqalulik formation quaternary cover undifferentiated cretaceous– paleocene deposits itilli formation ice sea/lake fault 84 the itilli formation is also exposed in the eastern part of the svartenhuk halvø area (fig. 73; j.g. larsen & pulvertaft 2000) where it was cored in the umiivik-1 well (dam et al. 1998b). on nuussuaq, the lower part of the formation (early campanian and older) is only present west of the kuugannguaq–qunnilik fault (chalmers et al. 1999 fold-out 2) where it is exposed in the itilli valley area. in this area, the formation was cored in the fp94-9-01 (fig. 65; madsen 2000) and gant #1 wells (fig. 74; dam 1996a). type section. the type section of the itilli formation is located in the southern part of the itillli valley along the pingunnguup kuua river and its tributaries, ukalersalik and anariatorfik (figs 65, 66, plate 2). the type section was described briefly by j.m. hansen (1976, 1980a) and in detail by dam & sønderholm (1994). neither the base nor the top of the formation is exposed in the type section. the base of the type section is located at 70°36.35´n, 54°13.79´w. reference sections. reference sections showing the lower boundary with the atane formation are exposed at kussinerujuk and asuk on the north coast of disko (figs 55, 77) and in the aaffarsuaq valley on central nuussuaq (fig. 83). the upper boundary towards the kangilia formation is located on the north coast of nuussuaq around kangilia (figs 72, 74, 87; birkelund 1965; rosenkrantz 1970; henderson et al. 1976). the upper boundary towards the eqalulik formation is exposed at qilakitsoq on central nuussuaq (figs 82, 83). well reference sections are available from the fully cored boreholes fp94-9-01 (madsen 2000), umiivik-1 (dam 1997), and gant#1 (dam 1996a). the formation was drilled but not cored in the gro#3 well (fig. 67; kristensen & dam 1997). for details on these wells, see below under description of members. thickness. the formation is more than 1400 m thick at the type locality where the base and top of the formation are not exposed (plate 2). in the nearby gro#3 well, the formation is more than 2000 m thick (fig. 67). the formation thins dramatically eastwards across the pebblessandmud not exposed (dyke) not exposed 550 500 470 700 650 600 800 m 750 fig. 66. expanded sedimentological log showing a representative portion of the itilli formation in the type section along the pingunnguup kuua river; the entire section, at a smaller scale, is shown in plate 2. for legend, see plate 1; for location, see fig. 65. from dam & sønderholm (1994). 85 kuugannguaq–qunnilik fault and is only 240 m thick along the northern slopes of the aaffarsuaq valley between qilakitsoq and tunoqqu. at kangilia, on the north coast of nuussuaq, the formation is at least 250 m thick (fig. 87), and the nearby gant#1 well penetrated 645 m of the formation without reaching the base of the formation. on svartenhuk halvø, the drilled and cored part of the formation is 960 m thick (excluding twenty-two paleocene dolerite intrusions with a total thickness of 240.2 m), but the base of the formation was not reached (fig. 75). at kussinerujuk on the north coast of disko, the measured part of the formation is 42 m thick (fig. 77; pulvertaft & chalmers 1990). lithology. in the type section, the itilli formation comprises mudstones, thinly interbedded sandstones and mudstones, chaotic beds, amalgamated beds of coarsegrained to very coarse-grained sandstone, and giant-scale cross-bedded sandstones (figs 66, 68, 69, 70, plate 2; dam & sønderholm 1994). these lithological contrasts are reflected in the blocky pattern in the petrophysical logs of the gro#3 well (fig. 67; kristensen & dam 1997). in the aaffarsuaq area, the itilli formation comprises mudstones, thinly interbedded sandstones and mudstones, and amalgamated sandstone and conglomerate units together with chaotic beds comparable to those in the itilli valley (dam et al. 2000). the main differences between the itilli and aaffarsuaq areas are that most of the channellised amalgamated sandstone and conglomerate units in aaffarsuaq are separated from the underlying deposits by major erosional surfaces or minor angular unconformities and that they are generally coarsergrained than their itilli counterparts. furthermore, the chaotic beds in aaffarsuaq occur at many levels in the section and are not restricted to a position immediately underlying a channellised sandstone unit, and the interbedded sandstone and mudstone units often show an overall thinning-upward trend. the amalgamated sandstone and conglomerate units of the aaffarsuaq area consist of very coarseto mediumgrained sandstone and conglomerate beds grading upward into thinly interbedded sandstones and mudstones. these coarse-grained units are up to 50 m thick, extend laterally beyond the extent of outcrop (several hundred metres) con. gr 100 m 0 300 40 200 300 400 500 600 700 800 900 1000 1100 1200 1300 1400 1500 1600 1700 1800 1900 2000 2100 2200 2300 2400 2500 2600 2700 2800 2900 3000 t.d.: 2996 m va ig at f m strat.stage gro#3 series a ga td al f m eq al ul ik f m k an gi lia f m it ill i f or m at io n a na ri ar to rf ik m em be r lo w er m aa st ri ch t. c am pa ni an up pe r m aa st ri ch t. pa le oc en e su bm ar in e ca ny on su bm ar in e ca ny on sl op e u pp er c re ta ce ou s n o pa ly no m or ph s re co rd ed d ue t o th er m al in flu en ce d an ia n hyaloclastite intrusion mudstone sandstone 240 dtc fig. 67. lithological log of the gro#3 well interpreted from petrophysical data. gr, spectral gamma-ray log; dtc, sonic log; con., coniacian; maastricht., maastrichtian; strat., lithostratigraphy; t.d., total depth. modified from christiansen et al. (1999). for location, see fig. 65. fig. 68. outcrop of the itilli formation in the type locality in the anariartorfik gorge showing amalgamated sandstones and thinly interbedded sandstones and mud stones (prominent sandstone unit is approxi mately 30 m thick). anariartorfik section 930–1010 m (plate 2). for location, see fig. 65. fig. 69. thinly interbedded sandstones and mudstones of the itilli formation in the anariartorfik section (1070–1150 m; plate 2). for location, see fig. 65. fig. 70. amalgamated sandstones of the itilli formation (anariartorfik member) deposited from turbidity currents in slope channels overlying contorted mudstones (at river level). pingunnguup kuua section 460–510 m (plate 2). the amalgamated sandstone unit is approximately 30 m thick and is cut by a dyke. for location, see fig. 65. 86 87 and include clasts of intraformational mudstone, sandstone (probably derived from the underlying atane for mation), redeposited concretions and basement litho l ogies. on the north coast of nuussuaq and on svartenhuk halvø, the formation is dominated by mudstones, thinly interbedded sandstones and mudstones, and bioturbated thinly interbedded sandstones and mudstones. these lithologies are arranged in coarsening-upward successions 10–50 m thick (fig. 75). at kussinerujuk on the north coast of disko, the formation consists of sandstones and conglomerates with angular cobbles and boulders of mudstone, sandstone, or interbedded sandstone and mudstone. other clasts consist of rounded pebbles of quartz and clay ironstone (fig. 77; pulvertaft & chalmers 1990, fig. 6). at asuk, the formation consists of a basal sandstone or conglomerate bed succeeded by sand-streaked mudstones (fig. 55). fossils. ammonites, belemnites, inoceramid bivalves (fig. 71) and rare crustaceans and corals are present in the formation at several localities on nuussuaq and svartenhuk halvø (birkelund 1965; rosenkrantz 1970; dam et al. 2000). however, most of these are not found in situ in the outcrops and may furthermore be reworked intraformationally or derived from older deposits. dinocysts, spores and pollen occur throughout in small numbers but cannot be determined neither in the lower part of the umiivik-1 and gro#3 wells nor from the outcrops in the itilli valley. burrows are occasionally present at the top of sandstone beds. details of fossil contents are presented under the description of individual members. depositional environment. the itilli formation was primarily deposited in slope and submarine fan environments. in the type area the amalgamated sandstones were deposited in slope channels initiated during sea level lowstands whereas the mudstones and the thinly interbedded sandstones and mudstones were deposited in interchannel slope areas (dam & sønderholm 1994). details on the depositional environment are presented under the description of individual members. boundaries. on the north coast of nuussuaq, the itilli formation overlies the atane formation in a small and poorly exposed section at high altitude immediately east of the ikorfat fault (fig. 22; a.k. pedersen et al. 2006b). the boundary with the underlying atane formation is well exposed on central nuussuaq along the northern slope of the aaffarsuaq valley between qilakitsoq and tunoqqu (fig. 82; nøhr-hansen 1996; dam et al. 2000) and on northern disko at kussinerujuk and asuk (fig. 2). in aaffarsuaq, the lower boundary is placed at the unconformity between the santonian deltaic deposits of the atane formation and the first turbidite sandstone or mudstone deposits of the campanian aaffarsuaq member (fig. 83). at kussinerujuk and asuk, the lower boundary is an erosional unconformity between the deltaic deposits of the atane formation and the slope deposits of the itilli formation (figs 55, 77, 78, 79). the upper boundary is well exposed on northern nuussuaq where it is defined at a major unconformity with the upper maastrichtian – lower paleocene anner tuneq conglomerate member of the kangilia formation (figs 72, 76, 87). however, in areas outside the distri bution of the annertuneq conglomerate member, there 28 cm fig. 71. giant inoceramid bivalves on top of a turbidite sandstone bed of the itilli formation (aaffarsuaq member) in the second ravine east of qilakitsoq. for location, see fig. 82. 88 is no lithological contrast at the boundary between the itilli and kangilia formations and in these areas bio stratigraphic data may be essential for the distinction between the turonian to lower maastrichtian itilli formation and the mainly danian kangilia formation. on svartenhuk halvø, the formation is overlain by hyaloclastite of the vaigat formation or, as observed on the east slope of firefjeld, by a conglomeratic unit that may be correlated with either the agatdal formation or with the quikavsak formation. the upper boundary of the itilli formation is not exposed on northern disko. on the north coast of nuussuaq and in the gant#1 and gro#3 wells, the itilli formation is unconformably overlain by the kangilia formation (figs 67, 72, 87). along the aaffarsuaq valley, the upper boundary is ambiguous and can be difficult to identify. east of qilakitsoq, however, campanian deposits of the itilli formation are overlain by a paleocene conglomerate/sandstone unit referred to the eqalulik formation (fig. 83) thus indicating a major unconformity. at nassaat, a tributary on the south-eastern side of agatdalen, the presence of upper santonian mudstones with spheno ceramus was reported by rosenkrantz (1970); this succession is now referred to the itilli formation. at this locality, the mudstones are overlain by pebbly sandstones and bituminous mudstones, a unit that referred to the agatdal formation by rosenkrantz (1970 fig. 4c) but to the upper atanikerdluk formation (quikavsak formation of this paper) by koch (1959 fig. 37). no new data are available from this outcrop. geological age. a ?late turonian to early maastrichtian age range for the formation is indicated by the ammonite fauna (fig. 16). palynomorph data are generally in accordance with the ammonite data, but local discrepancies occur, probably due to redeposition of ammonites from the underlying deposits. at asuk and kussinerujuk on northern disko, the itilli formation is of cenomanian–turonian age (fig. 16; bojesen-koefoed et al. 2007) whereas on svartenhuk halvø, the formation is of turonian to early campanian age (fig. 16; nøhr-hansen 1996, unpublished data; dam et al. 1998b). strata of campanian age referred to the hilli formation are also exposed on nuussuaq (itilli, aaffarsuaq, agatdalen; birkelund 1965; sønderholm et al. 2003). the maastrichtian part of the formation is only known from kangilia and aaffarsuaq (birkelund 1965; rosenkrantz 1970; nøhr-hansen 1996) and has been dated in the gro#3 well (sønderholm et al. 2003). detailed information on the age of the various members is found in the description of the members below. correlation. the pre-early campanian part of the formation is coeval with the atane and upernivik næs for mations (fig. 16). subdivision. four members are recognised, the anariartorfik, umiivik, kussinerujuk, and aaffarsuaq members, of which the anariartorfik member is found only west of the kuugannguaq–qunnilik fault. the anariartorfik member consists of interbedded turbidite channel sandstones and mudstones deposited in a faultcontrolled slope environment. the umiivik member is the northernmost representative of the itilli formation. it is finer grained than the anariartorfik member, and dominated by major coarsening-upward successions it ill i fo rm at io n k an gi lia fo rm at io n fig. 72. unconformity between the itilli formation (umiivik member) and the overlying kangilia formation with the resistant pale-weathering annertuneq conglomerate member at the base. north coast of nuussuaq, at annertuneq. base of conglomerate is at c. 300 m a.s.l. for location, see fig. 74. 89 deposited in a base-of-slope and basin-floor fan environment. both the anariartorfik and umiivik members record deposition from probably the cenomanian to the maastrichtian. the kussinerujuk member is only exposed on northern disko at asuk and kussinerujuk and represents a cenomanian transgressive event on top of the atane formation. the aaffarsuaq member crops out east of the kuugannguaq–qunnilik fault on central nuussuaq and comprises a relatively thin wedge of lower to middle campanian strata that are distinctly different from the anariartorfik member with regard to their large content of intraformational clasts, the overall lack of fine-grained mudstones and the lenticular shapes of the thinly interbedded sandstone and mudstone units. the sediments of the aaffarsuaq member were deposited in a deep marine, channellised footwall fan system. anariartorfik member new member history. strata referred here to the anariartorfik member were described informally as the itilli formation by j.m. hansen (1980b). detailed studies of strata assigned to the anariartorfik member were published by dam & sønderholm (1994) and madsen (2000). name. the member is named after the anariartorfik valley leading into the itilli valley (fig. 65). distribution. the anariartorfik member is exposed in the itilli valley and is present in the subsurface west of the kuuganguaq–qunnilik fault in the western part of nuussuaq (fig. 11). type section. the type section is the same as for the itilli formation in the southern part of the itilli valley (figs 65, 66, plate 2). the base of the type section is located at 70°36.35´n, 54°13.79´w. reference sections. well-exposed reference sections are located in the northern part of the itilli valley and in the adjoining tunorsuaq valley (figs 2, 65, 74). the member was cored in the fp94-9-01 borehole between 32 and 522 m in the central part of the itilli valley (fig. 65; madsen 2000); the cores are stored at geus in copenhagen. the member was drilled but not cored in the gro#3 well between 959 m and 2996 m (figs 65, 67; kristensen & dam 1997). thickness. the member has a thickness (including intrusions) of at least 2037 m in the gro#3 well. approxi mately 1400 m are exposed in two incomplete sections in the itilli valley (fig. 65, plate 2). lithology. in the type section, the anariartorfik member is dominated by mudstones and thinly interbedded sandstones and mudstones, chaotic beds, amalgamated beds of coarse-grained to very coarse-grained sandstones, and giant-scale cross-bedded sandstones (figs 66, 68, 69, 70, plate 2; dam & sønderholm 1994). these lithological contrasts create a characteristic blocky pattern in the petrophysical log of the gro#3 well (fig. 67; kristensen & dam 1997). the mudstones are dark grey to black, show parallel lamination and occur in intervals up to 15 m thick in the type section. persistent layers of early diagenetic ankerite concretions occur in most mudstone intervals. the thinly interbedded sandstone and mudstone units comprise laterally extensive graded laminae and beds of fineto very coarse-grained sandstone capped by parallel-laminated mudstones; this facies forms successions several tens of metres in thickness (plate 2). the sandstone beds range in thickness from less than 1 cm to 80 cm, have flat, locally scoured bases and show normal grading; the beds may be structureless near their base or may be parallelor cross-laminated throughout. the sand stone beds are laterally persistent at outcrop scale (> 100 m) and there is generally no systematic variation in thickness. convolute bedding and slump structures are common. early diagenetic ankerite concretions are common in the mudstone units. the chaotic beds are up to 30 m thick and consist of contorted, laminated mudstone, thinly interbedded sandstone and mudstone, and homogenised mudstone with scattered sand grains, reworked early diagenetic ankerite concretions, semi-indurated mudstone and sandstone clasts and occasional basement clasts. the chaotic beds invariably underlie a thick unit of amalgamated sandstone beds and are often associated with sandstone dykes (fig. 66, plate 2). the amalgamated sandstone beds form up to 50 m thick channel-shaped bodies that can be followed for 1–2 km along strike (figs 66, 68, 70, plate 2). the sandstone units are erosionally based and locally show welldeveloped flute casts and channel-shaped scours at the base. the sandstone units may show a thinning-upward trend and have a gradational or sharp, but non-erosional contact to the overlying interbedded sandstone and mudstone units. internally, the sandstone units are dominated by amalgamated, normally graded, medium-grained 90 to pebbly, very coarse-grained sandstone beds (0.1–3.5 m thick) or, less commonly, show planar and trough cross-bedding. the graded beds have planar, erosional bases and can be followed laterally for more than 140 m without major variations in thickness (figs 66, 68, 70). angular mudstone clasts, ranging from a few millimetres to 45 cm across and rounded basement pebbles up to 8 cm across are common in the graded beds. in some cases, the uppermost part of the graded sandstone beds show well-developed parallel lamination, with parting lineation and tool marks, trough cross-bedding, low-angle cross-bedding and climbing ripple lamination. the graded sandstone beds are commonly separated by thin mudstone beds or units of thinly interbedded mudstone and sandstone. giant-scale cross-bedded sandstone units occur as channel-shaped bodies at a few levels in the type section. the units comprise 7–15 m thick low-angle, crossbedded sets dominated by coarseto medium-grained sandstone; the sets can be followed for approximately 150 m in a dip direction. at one locality, a single set fills out a large channel-shaped erosional depression that is approximately 300 m wide (plate 2 at 800 m; dam & sønderholm 1994). fossils. macrofossils have not been found in this member, but burrows are locally present at the top of the sandstone beds. they include common ophiomorpha isp., thalassinoides isp., escape burrows, and rare helmin thopsis isp. (dam & sønderholm 1994). identifiable palynomorphs are very rare in the area of the type sec66 ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ itsaku umiiviup kangerlua kangiusap imaa fi re fje ld uparuaqqusuitsut qeqertarsuaq karrat fjord intrusion kangilia formation precambrian basement ice sea/lake upernivik næs formation vaigat formation svartenhuk formation quaternary cover quikavsak/agatdal formation itilli formation umiivik member fig. 37 fig. 36 eqalulik formation 10 km 54° 71°30' ■ ■ ■ ■ ■ normal fault (tick on downthrow side) major fault with variable sense of movement through time monoclinal flexure drill hole ■ 400709 400711 400710 400712 400708 umiivik-1 fig. 73. geological map of the south-eastern part of svartenhuk halvø and the west coast of qeqertarsuaq showing the location of the umiivik-1 well (type section of the umiivik member of the itilli formation) and the position of shallow wells drilled by ggu in 1992 (ggu 400708–400712). reference sections of the upernivik næs formation on qeqertarsuaq are indicated (see figs 36, 37). for location, see fig. 2. modified from j.g. larsen & grocott (1991) and j.g. larsen & pulvertaft (2000). 91 v ai ga t fo rm at io n eq al ul ik f or m at io n q ua te rn ar y co ve r it ill i f or m at io n, u m iiv ik m em be r st ri ke a nd d ip o f s tr at a d ri ll ho le a ba nd on ed s et tle m en t k an gi lia f or m at io n w ith a c m em be r in tr us io n ic e se a/ la ke 22 18 40 07 07 40 07 06 fp 94 -1 102 fp 94 -1 104 40 07 05 u i n g a j a a r s u a q 3 6 5 25 00 m d r v k tu no rs ua q 53 °3 0’ 53 °1 5’ 70 °4 5’ d an ien klø ft tu no rs ua q n ia qo rn at tu pe rs ua rt aa it e rl a a n ia qo rs ua q a nn er tu ne q k an gi lia h am ite sk lø ft sa vi aq qa t fp 94 -1 105 se rf at g a n t # 1 12 90 fi g. 7 4. g eo lo gi ca l m ap o f t he n or th c oa st o f n uu ss ua q sh ow in g ou tc ro ps o f t he i ti lli a nd k an gi lia f or m at io ns . t he w el ls f p9 411 -0 2, 04 a nd 05 a re lo ca te d ne ar s er fa t. t he sh al lo w w el ls d ri lle d by g g u in 1 99 2 (g g u 4 00 70 5– 40 07 07 ) ar e lo ca te d at a nn er tu ne q. t he g a n t #1 w el l i s in th e tu no rs ua q va lle y. d r , d an ie nr yg ge ; a c , a nn er tu ne q c on gl om er at e m em be r. c on to ur in te rv al is 2 00 m . f or lo ca ti on , s ee f ig . 2 . m od ifi ed fr om r os en kr an tz e t a l. 19 74 . 92 tion, but are abundant in the cuttings from the middle and upper part of the gro#3 well (nøhr-hansen 1997c). depositional environment. deposition of the anariartorfik member took place in a fault-controlled slope environment (dam & sønderholm 1994). most of the channellised amalgamated turbidite sandstone beds were deposited from high-density turbidity currents in confined low-sinuosity channels, although a few examples of giant-scale, cross-bedded sandstones are interpreted as the products of lateral accretion in meandering slope channels. the thinly interbedded sandstones and mudstones are interpreted as having been deposited from traction currents and from fall-out processes associated with various sedimentation stages within waning low-density currents. the lateral continuity and the lack of systematic vertical thickness variations in the thinly interbedded sandstones and mudstones suggest that these deposits were not confined by channel-levee systems, but most likely represent interchannel slope deposits. where present, the chaotic beds always underlie undisturbed channel sandstones. this suggests that the channels were initially excavated by retrogressive slumping of unstable sediments on the slope followed by channel excavation by scouring (dam & sønderholm 1994). boundaries. the lower and upper boundaries of the anariartorfik member are not exposed. the upper boundary was drilled in the gro#3 well but not cored (fig. 67). an erosional unconformity between the anariartorfik member and the overlying kangilia formation was suggested by kristensen & dam (1997). geological age. palynomorphs in the gro#3 well indicate a ?coniacian/santonian – early maastrichtian age for the uppermost c. 525 m of the anariartorfik member. in the lowermost 1510 m of the well, the organic material is degraded due to thermal maturation and this part of the formation cannot be dated (nøhr-hansen 1997c). umiivik member new member history. outcrops of marine upper cretaceous strata referred to the umiivik member on svartenhuk halvø and northern nuussuaq have only been briefly described by earlier workers since their main focus was on the collection of fossils (rosenkrantz et al. 1942; birkelund 1956, 1965; rosenkrantz 1970). later studies in connection with mapping by the geological survey and drilling of stratigraphic wells in the umiivik area of svartenhuk halvø have been reported by christiansen (1993), dam et al. (1998b), christiansen et al. (2000) and j.g. larsen & pulvertaft (2000). on northern nuussuaq, extensive stratigraphic studies have been carried out at annertuneq and kangilia (christiansen et al. 1992; christiansen 1993; christiansen et al. 1994) and at serfat in connection with drilling of mineral exploration wells (dam & nøhr-hansen 1995). name. the member is named after the stretch of shore located south-east of firefjeld in the inner part of the umiiviup kangerlua bay, eastern svartenhuk halvø (fig. 73). distribution. on svartenhuk halvø, the umiivik member is locally exposed in the north-western part of the peninsula (fig. 2), and more extensively in the low-lying coastal outcrops around the bay of umiiviup kangerlua, and c. 5 km north-west of the south-eastern corner of the peninsula (fig. 73). on nuussuaq, exposures are found along the north coast between niaqornat and ikorfat and in the tunorsuaq valley (figs 2, 74). on svartenhuk halvø, the member has been cored in five shallow wells (ggu 400708–712; christiansen 1993; christiansen et al. 1994a) and in the deep umiivik1 well (fig. 73; christiansen et al. 1994; nøhr-hansen 1996; christiansen et al. 1997; dam 1997; nøhr-hansen 1997a; dam et al. 1998b). on the north coast of nuussuaq the member has been penetrated in six shallow boreholes: ggu 400705–707 and fp4-11-02, fp94-11-04 and fp94-11-05 (chri stiansen et al. 1994; dam & nøhr-hansen 1995), and in the deep gant#1 well (fig. 74; christiansen et al. 1996c; dam 1996a; dahl et al. 1997; nøhr-hansen 1997b; kierkegaard 1998). type section. the most complete section of the member is available in the umiivik-1 core drilled on the southern shore of umiiviup kangerlua on svartenhuk halvø (ggu 439301; figs 73, 75). neither the base nor the top of the member is, however, seen in this well. the cores are stored at geus in copenhagen. the umiivik-1 well is located at 71°36.70´n, 54°02.52´w. reference sections. the best-exposed outcrops of the umiivik member are found below the kangilia formation (annertuneq conglomerate member) at annertuneq and kangilia on the north coast of nuussuaq (figs 72, 93 87). at annertuneq, the ggu cores 400705–707 were drilled to supplement the outcrop studies (fig. 74; christiansen et al. 1994). in the gant#1 well, the interval from 255.8 m to 900.6 m (fig. 76) is assigned to the umiivik member (dam 1996a). the cores are stored at geus in copenhagen. thickness. the total thickness of the umiivik member is unknown, but in the umiivik-1 well on svartenhuk halvø it is at least 960 m thick (excluding a total of 22 dolerite intrusions with a cumulative thickness of 240.2 m). in the slopes behind the umiivik-1 drill site, approximately 185 m of poorly exposed mudstone is present below the base of the hyaloclastic volcanic rocks of the vaigat formation – this part of the succession is partly penetrated by the cores 400710–711 (fig. 73; christiansen et al. 1994). in the gant#1 well, on northern nuussuaq, 620 m have been cored (excluding intrusions). lithology. the umiivik member is dominated by black mudstones intercalated with laminae and thin beds of sandstone; carbonate concretions are common (figs 75, 900 n o pa ly no m or ph s re co rd ed d ue t o th er m al in flu en ce fr om ig ne ou s in tr us io ns vfsiclay siclay siclayvc sand f m c 1200 1100 1000 vf vc sand f m c vf vc sand f m c ★ ★ ★ ★ ★ ★ ★ ★ ★ u pp er c on ia ci an u pp er t ur on ia n – ?u pp er c on ia ci an u pp er t ur on ia n 200 300 400 m 100 overburden ★ ★ ★ ★ ★ ★ ★ ★ umiivik-1 700 800 ?u pp er t ur on ia n 500 600 ★ ★ ★ ★ fig. 75. sedimentological log showing type section of the umiivik member (itilli formation) in the umiivik-1 well. for location, see fig. 73; for legend, see plate 1. modified from dam et al. (1998b). 94 700 650 610 vfsiclay f fm c m cvc pebblessand vfsiclay f fm c m cvc pebblessand vfsiclay f fm c m cvc pebblessand 450 400 350 310 600 550 500 150 100 50 m overburden 300 250 200 gant#1 900 850 800 750 it ill i f or m at io n u m iiv ik m em be r it ill i f or m at io n u m iiv ik m em be r it ill i f or m at io n u m iiv ik m em be r a nn er tu ne q c on gl om er at e m em be r k an gi lia f or m at io n 95 76). heavily bioturbated interbedded sandstones and mudstones, chaotic beds, and structureless, muddy sandstones also occur. indistinct thickeningand coarseningupward cycles, 2–85 m thick, can be seen in the umii vik-1, fp94-11-02 and fp94-11-04 wells (fig. 75; dam & nøhr-hansen 1995). a cobble conglomerate composed of lithified sandstone clasts in a mudstone matrix is exposed in stream gullies on the south side of uparu aqqusuitsut on north-eastern svartenhuk halvø (fig. 73; christiansen et al. 2000). in the gant#1 core, chaotic beds, muddy sandstones and sandy mudstones alternate with thick, sharpbased fining-upward successions (fig. 76; dam 1996a). the fining-upward successions consist of amalgamated sandstone beds grading upward into thinly interbedded sandstones and mudstones. thin coarsening-upward successions also occur. the amalgamated sandstones consist of normally graded, mediumto coarse-grained sandstone beds with scattered basement pebbles and mudstone intraclasts. the sandstone beds are generally structureless, but parallel and cross-lamination occurs towards the top of some beds. a thin mudstone layer usually caps the sandstone beds. the thinly interbedded sandstones and mudstones consist of sharp-based, graded laminae and beds of finegrained to very coarse-grained sandstone alternating with black parallel-laminated mudstones. the sandstones are well-sorted and may show parallel and cross-lamination; small mudstone rip-up clasts frequently occur throughout the sandstone beds. fossils. ammonites from the umiivik member have been described from several localities on svartenhuk halvø and include scaphites mariasensis umivikensis (birke lund), scaphites preventricosus svartenhukensis (birke lund), clioscaphites sp. aff. saxitonianus (mclearn), scaphites cobbani (birkelund), scaphites rosenkrantzi (birkelund), scaphites cf. corvensis (cobban), clioscaphites saxitonianus septentrionalis (birkelund), ammonites of the genus haresiceras and inoceramids of the steenstrupi group (birkelund 1965). belemnites from svartenhuk halvø include actinocamax cf. primus (arkhangelsky) and actinocamax sp. (birkelund 1956). unidentified teleost fish remains have also been found in strata yielding coniacian ammonites (bendix-almgreen 1969). ammonites occur at several localities along the north coast of nuussuaq in the umiivik member, including pseudophyllites skoui (birkelund), scaphites (hoploscaphites), s. (h.) greenlandicus (donovan), s. (h.) ravni (birkelund), and s. (h.) ikorfatensis (birkelund). the belemnites actinocamax groenlandicus (birkelund) and actinocamax aff. groenlandicus (birkelund), and an indeterminate solitary corallum have also been collected (birkelund 1956, 1965; floris 1972). dinocysts are abundant in the umiivik member (dam & nøhr-hansen 1995; nøhr-hansen 1996, 1997a; dam et al. 1998b). depositional environment. the umiivik member records deposition from low-density and high-density turbidity currents, debris flows, slumping and fall-out from suspension. deposition of the mudstones and intercalated mudstones and sandstones took place in a base-of-slope and basin-floor fan environment. the succession in the gant#1 well, which is situated close to the kuugan guaq–qunnilik fault, reflects a fault-controlled base-ofslope environment with major and minor distributary feeder channels, small turbidite lobes and interdistributary channel areas. boundaries. the umiivik member unconformably overlies the atane formation on the north coast of nuussuaq in a small and poorly exposed section at high altitude immediately east of the ikorfat fault (fig. 22; a.k. pedersen et al. 2006b). west of the ikorfat fault, the lower boundary of the umiivik member is not exposed and has not been drilled. on northern nuussuaq, the member is unconformably overlain by the kangilia formation (figs 72, 74, 87). at most localities on svartenhuk halvø, it is unconformably overlain by paleocene hyaloclastic rocks of the vaigat formation (fig. 73). at firefjeld, however, a thin paleo cene conglomeratic unit that may be correlated with either the agatdal formation or the quikavsak formation is present between the umiivik member and the volcanic succession. geological age. the age of the umiivik member is based on dinocyst and ammonite data (birkelund 1965; nøhrhansen 1996, 1997a). the ammonites from svartenhuk halvø indicate a ?late turonian – early campanian age for the umiivik member in this area (fig. 16). dinocysts facing page: fig. 76. sedimentological log of the umiivik member (itilli formation) and the annertuneq conglomerate member (kangilia formation) in the gant#1 well. for location, see fig. 74; for legend, see plate 1. modified from dam (1996a). 96 from the umiivik-1 well indicate a ?late turonian – late coniacian age for the uppermost 659 m of the well; below this palynomorphs are not preserved due to severe alteration of the mudstones by the paleocene dolerite intrusions. it is likely, however, that the core includes cenomanian–turonian strata (nøhr-hansen 1997a; dam et al. 1998b). dinocysts from outcrops and from the three shallow wells (ggu 400710, 400711, 400712) along the southern shore of umiiviup kangerlua on svartenhuk halvø indicate a santonian to early cam panian age (nøhr-hansen 1996). on the north coast of nuussuaq, the ammonites indicate the presence of upper campanian and maastrichtian strata which is in accordance with palynostratigraphic dating (birkelund 1965; nøhr-hansen 1996). kussinerujuk member new member history. on northern disko at kussinerujuk and asuk, local outcrops of conglomerate, sandstone and mudstone unconformably overlying the atane formation have previously been correlated with the paleocene kangilia formation (j.m. hansen 1980b; pulvertaft & chalmers 1990). later, these beds were included in the asuup innartaa member of the danian quikavsak formation (dam 2002). on the basis of new biostratigraphic data, however, these are here assigned to the kussinerujuk member of cenomanian–turonian age (bojesen-koefoed et al. 2007). name. the member is named after kussinerujuk on the north coast of disko, where it is exposed in narrow gorges (fig. 2). distribution. the member is only known from the north coast of disko around kussinerujuk and asuk. type section. the type section of the member is in the ravines at kussinerujuk (figs 2, 77, 78). the type section is located at 70°13.10´n, 53°27.68´w. 0 10 20 30 40 poorly exposed a ta ne f or m at io n k in gi tt oq m em be r it ill i f or m at io n k us si ne ru ju k m em be r solifluction, volcanic scree 50 m clay silt sand pebbles, cobbles vf f m c vc fig. 77. type section of the kussinerujuk member (itilli formation) at kussinerujuk, north coast of disko. for location, see fig. 2; for legend, see plate 1. base of section is c. 450 m a.s.l. 97 reference sections. reference sections are found in the coastal cliffs east of asuk (figs 55, 79). the lithologies here differ from those of the type section but are locally affected by complicated deformation resulting from landslides obscuring the stratigraphic relationships (bojesenkoefoed et al. 2007). thickness. at kussinerujuk the measured part of the formation is at least 42 m thick (fig. 77; see also pulvertaft & chalmers 1990) but the member may be up to a couple of hundred metres thick. lithology. the kussinerujuk member comprises thinly interbedded mudstones and sandstones, and mudstone clast conglomerates, first described by pulvertaft & chalmers (1990 fig. 6). at kussinerujuk the member consists of mud-clast conglomerates of varying thicknesses, with angular clasts of cobble to boulder size, interbedded with mediumto coarse-grained sandstone (fig. 77). in some conglomerate beds, the clasts are imbricated whereas other conglomerates appear to be disorganised. pebble-sized clasts consist of rounded quartz and clay ironstone while angular cobbles and boulders consist of mudstone, sandstone, or interbedded sandstone and mudstone. at asuk, the atane formation is overlain by mudstones and sandstones that are referred to the kussinerujuk member. these sediments occur in three separate outcrops. the westernmost of these is seen in fig. 79, the central outcrop is seen in fig. 80, and the eastern is seen in the coastal cliff adjacent to the huge alluvial fan at qorlortorsuaq. the western outcrop is overlain by volcanic breccias and lavas of the asuk member (lower part of the vaigat formation; bojesen-koefoed et al. 2007). the volcanic rocks are part of a major landslide, and the complicated relationships seen in fig. 79 suggest that the sediments at this locality are also affected by sliding; their stratigraphic relationships with the central and eastern outcrops are thus uncertain. a simplified sedimentological log from the central outcrop is shown in fig. 55. here a basal conglomerate of mudstone and sandstone clasts in a matrix of coarsegrained sand erosionally overlies the atane formation and is succeeded by parallel-laminated sandstones interbedded with sand-streaked mudstones (figs 55, 80). the western, landslipped outcrops comprise a unit of parallel-laminated, sand-streaked mudstone, locally with small-scale load structures (fig. 81a). other units consist of strongly erosional, channellised sandstones that cut into mudstones (fig. 81 b); thicker sandstone beds are fineto medium-grained and show parallel lamination and ripple cross-lamination. other sandstone beds are structureless and have erosional bases with sole-marks. fossils. brackish to marine dinocysts occur in the member and are relatively abundant at asuk (mcintyre 1994a, b) noted that rich pollen and spore assemblages dominated by bisaccate conifer pollen are present in the type section whereas dinocysts are very rare. depositional environment. the kussinerujuk member records a transgressive event on top of a variably incised atane fm 3 m kingittoq mb kussinerujuk mb itilli fm fig. 78. type locality of the kussinerujuk member (itilli formation) at kussinerujuk. note the basal erosional boundary with the atane formation beneath; the uppermost depositional unit underlying the kingittoq member is a dark grey, delta front mudstone. the kussinerujuk member is dominated by mud-clast conglomerates with a matrix of sand (see fig. 77). for location, see fig. 2. 98 surface capping the atane formation east of the kuuganguaq–qunnilik fault. the genetic relationships between the two outcrop areas of the member are, however, poorly understood. the conglomerates at the type locality at kussinerujuk were deposited from channellised sediment gravity flows at unknown water depths possibly generated by slope failure in a deeply incised channel. farther east, at asuk and qorlortorsuaq, the mudstonedominated succession is tentatively interpreted to be deposited in a marine environment during transgression. deposition occurred below storm wave base on a gently sloping sea floor from low-energy unconfined turbidity currents or distal storm-generated flows. boundaries. the lower boundary, as exposed at kussinerujuk, asuk and qorlortorsuaq (fig. 2), is an erosional unconformity where mudstones or mud-clast conglomerates sharply overlie deltaic sandstones of the atane formation (kingittoq member; figs 16, 55, 77). the upper boundary is not exposed, but is probably an erosional unconformity overlain by paleocene hyaloclastite breccias of the vaigat formation. geological age. spores, pollen and dinocysts indicate a cenomanian to late turonian age for the member at kussinerujuk (mcintyre 1994a, b; this study). dinocysts indicate a cenomanian age at asuk (fig. 16; bojesenkoefoed et al. 2007). atane formation kingittoq member vaigat formation asuk member itilli formation kussinerujuk member e w fig. 79. outcrops of the itilli formation at the western end of the coastal cliff section at asuk. outcrops of the atane formation are seen farthest to the east (left); for detail, see fig. 57. the pale sandstones (centre-left) are impregnated with hydrocarbons (see bojesen-koefoed et al. 2007).the dark grey mudstones in the centre of the photograph are referred to the kussinerujuk member as are the sandstones to the west. the dark-weathering rocks cropping out above these sandstones are referred to the volcanic asuk member (vaigat formation). these volcanic rocks are part of a major landslide, and the complicated structural relationships between the pale sandstones and the mudstones of the itilli formation suggest that the coastal exposures may be affected by landslides. the height of the cliff is c. 30 m. for location, see fig. 2. atane fmitilli fm kingittoq mbkussinerujuk mb fig. 80. erosional unconformity between the atane formation and the itilli forma tion (see fig. 55) in the coastal cliff east of asuk. the coastal cliff is c. 30 high. for location, see fig. 2. 99 correlation. the kussinerujuk member is coeval with parts of the umiivik and anariartorfik members of the itilli formation. aaffarsuaq member new member history. the un-named upper cretaceous marine sandstones and shales exposed on central nuussuaq that were informally referred to the aaffarsuaq member by dam et al. (2000), are mainly known for the remarkable discoveries in 1952 of huge specimens of upper santonian – lower campanian inoceramid (sphenoceramus) bivalve shells (fig. 71; rosenkrantz 1970). name. the member is named after the aaffarsuaq valley on central nuussuaq (fig. 82). distribution. the member is found on central nuussuaq where it is well exposed along the south-facing slopes of aaffarsuaq between qilakitsoq and tunoqqu. minor exposures are present in a stream section on the northfacing slope of aaffarsuaq, south of nalluarissat, in kangersooq (fig. 82) and in turritelladal at scaphites næsen, and in agatdalen where the teltbæk fault crosses the agatdalen river (figs 82, 113; dam et al. 2000). type section. the type section is located in the second ravine east of qilakitsoq on the northern slopes of aaffarsuaq (figs 82, 83). the base of the type section is located at 70°28.70´n, 53°21.97´w. reference sections. well-exposed reference sections occur in ravines on the northern side of aaffarsuaq along nalluarissat, between qilakitsoq and tunoqqu (fig. 82). thickness. in the main outcrop area, the thickness of the member decreases eastward from c. 250 m in the type section to 65 m at tunoqqu. lithology. the aaffarsuaq member is characterised by amalgamated sandstone and rip-up mudstone and sandstone clast conglomerate units alternating with thinly interbedded sandstones and mudstones, dark, sandstreaked mudstones, and chaotic beds of homogeneous mudstone commonly cut by sandstone dykes and synsedimentary faults (figs 83, 84, 85, 86; dam et al. 2000). the amalgamated sandstone and conglomerate units of the aaffarsuaq member consist of very coarseto medium-grained sandstone and conglomerate beds grading upwards into thinly interbedded sandstones and mudstones. these coarse-grained units are up to 50 m thick and extend laterally beyond the extent of outcrop (several hundred metres) or occasionally form lenticular bodies. individual beds have planar erosional bases and internally show normal grading. beds are from <25 cm kussinerujuk member itilli formation a b fig. 81. a: silt-streaked mudstone with small-scale load structures in the kussinerujuk member (itilli formation) at asuk. lens cap for scale. b: land-slipped deposits of erosionally based sandstones interbedded with mudstones at asuk are assigned to the kussinerujuk member of the itilli formation. encircled persons for scale; for location, see fig. 2. 100 aga td al en 53 °2 5' 53 °1 5' 53 °0 5' 53 °0 0' 73 °3 0' 70 °3 2' 70 °3 1' tu no qq u itivnera 53 °2 0' 53 °1 0' 50 0 50 0 50 0 70 0 90 0 80 0 60 0 50 0 40 0 40 0 400 40 0 2 km 12 17 79 8 10 20 13 27 24 21 8 5 ib 8 30 ravine 3 ravine 4 ravine 2 ravine 1 qilakitsoq a af fa rs ua q n al lu ar iss at k an ge rs o o q v ai ga t fo rm at io n eq al ul ik f or m at io n a br ah am m em be r q ua te rn ar y co ve r it ill i f or m at io n, a af fa rs ua q m em be r st ri ke a nd d ip o f s tr at a c on to ur in m et re s a ta ne f or m at io n q ila ki ts oq m em be r ic e se a/ la ke 30 0 20 0 30 040 050 0 60 0 70 0 80 0 16 00 16 00 14 00 14 00 10 00 fi g. 8 2. g eo lo gi ca l m ap o f t he n or th si de o f t he a af fa rs ua q va lle y be tw ee n q ila ki ts oq a nd a ga td al en . t he ty pe se ct io n of th e q ila ki ts oq m em be r ( a ta ne f or m at io n) is lo ca te d al on g q ila ki ts oq , an d th e ty pe s ec ti on o f t he a af fa rs ua q m em be r (i ti lli f or m at io n) is lo ca te d in r av in e 2. i b , t yp e se ct io n of th e it iv ne ra b ed . c on to ur in te rv al is 1 00 m in th e se di m en ta ry o ut cr op s, 2 00 m in th e v ai ga t f or m at io n. m od ifi ed fr om r os en kr an tz e t a l. (1 97 4) . 101 160 140 120 280 260 240 220 200 180 not exposed volcaniclastic sediments not exposed 320 m 300 pebblessandmud vf f m c vc f m c pebblessandmud vf f m c vc f m c a af fa rs ua q m em be r a ta ne f m q ila ki ts oq m b it ill i f or m at io n va ig at f or m at io n eq al ul ik f m a br ah am m b it ill i f or m at io n a af fa rs ua q m em be r 100 80 60 40 20 0 fig. 83.type section of the aaffarsuaq member (itilli formation), second ravine east of qilakitsoq, compare with fig. 85. for location, see fig. 82; for legend, see plate 1. modified from dam et al. (2000). 102 to more than 5 m thick and are commonly separated by thin mudstone beds (figs 83, 85). the sandstones are mostly structureless, whereas the conglomerates show considerable lateral variation, are poorly sorted, and always have a coarse-grained sandy matrix. both matrixand clast-supported conglomerates occur. the clasts consist of intraformational mudstone, sandstone (probably derived from the underlying atane formation), concretions and basement lithologies (dam et al. 2000). the amalgamated sandstone and conglomerate units are overlain by thinly interbedded sandstone and mudstone units with either a gradational or a sharp contact. these units are up to 30 m thick and show a general thinningand fining-upward trend (figs 83, 85). they consist of laterally persistent, normally-graded, <1–40 cm thick beds of fineto medium-grained sandstone capped by parallel-laminated mudstone. mudstones with thin (<2 cm), laterally persistent sandstone streaks form monotonous successions 10–30 m thick. the mudstones are finely laminated and laminae are usually normally graded. the sandstone laminae are normally graded, parallelor cross-laminated throughout and constitute less than 25% of the succession. up to 15 m thick units of strongly contorted, thinly interbedded sandstones and mudstones and chaotic beds of homogeneous mudstone with evenly scattered coarse sand grains, mudstone intraclasts, transported early diagenetic concretions and basement clasts occur both within the mudstone-dominated successions and directly beneath the thick units of amalgamated sandstones and conglomerate. fossils. ammonites have been found at scaphitesnæsen in turritellakløft and include pseudophyllites skoui (birke lund), baculites obtusus (meek), baculites cf. haresi (reeside), scaphites cobbani (birkelund) and scaphites rosenkrantzi (birkelund) (birkelund 1965). reworked ammonites including scaphites cf. svartenhukensis (birke lund 1965) have been found at tunoqqu (dam et al. 2000). well-preserved inoceramids are commonly found at the top of turbidite sandstone beds, suggesting only little transportation (fig. 71). bryozoans fixed to inoceramid bivalve shells have been reported. a single, nearly complete capitulum of the cirripid eskimolepas gregersi rosenkrantz (rosenkrantz 1970) has also been found. dinocysts and pollen belonging to the aquilapollenites interval have been reported from scaphitesnæsen and tunoqqu (nøhr-hansen 1996; dam et al. 2000). depositional environment. the aaffarsuaq member exhibits a characteristic suite of thick units of amalgamated sandstone and conglomerate beds, interpreted to represent deposition mainly from gravity flows in a channellised, footwall fan system. deposition of the amalgamated sands was confined to major turbidite channels. the intervening thinly interbedded sandstone and mudstone units were deposited mainly from low-density turbidity currents confined to minor channels. the sand-streaked, mudstone-dominated units were deposited by waning lowdensity turbidity currents in an interchannel slope setting. the contorted, thinly interbedded sandstones and mudstones probably formed by downslope displacement of semi-consolidated sediment (dam et al. 2000). in the aaffarsuaq member, the chaotic beds are not invariably associated with turbidite channel deposits, as a ta ne f m q ila ki ts oq m b it ill i f m a af fa rs ua q m b fig. 84. unconformity between santonian deltaic deposits of the atane formation (qilakitsoq member) and campanian turbidite sandstones and conglomerates of the aaffarsuaq member (itilli formation). east-side of tunoqqu. for location, see fig. 82. 103 observed in the anariartorfik member, suggesting that downslope failure could also be associated with seismic activity (dam et al. 2000). geological age. the ammonites at scaphitesnæsen indicate an early campanian age (baculites obtusus zone) for the aaffarsuaq member. the mudstones are generally almost barren of palynomorphs but a rich flora is locally fig. 86. conglomerate at the top of the type section of the aaffarsuaq member (itilli formation); note the predominance of sedimentary clasts. ruler is 120 cm long. for location, see fig. 82. fig. 85. mudstones and conglomeratic sandstones of the upper part of the aaffarsuaq member (itilli formation) in the type section in the second ravine east of qilakitsoq, see fig. 83, c. 175–300 m. for location, see fig. 82. 104 present that suggest an early–middle campanian age for the member (nøhr-hansen 1996; dam et al. 2000). boundaries. the lower boundary is an angular unconformity. tilted strata of the deltaic qilakitsoq member (atane formation) were truncated during the transgression recorded by the aaffarsuaq member (dam et al. 2000). along the aaffarsuaq valley, the aaffarsuaq member is unconformably overlain by paleocene hyaloclastites of the vaigat formation (a.k. pedersen et al. 2002). however, in two ravines along aaffarsuaq, it is overlain by paleocene mudstones and volcaniclastic sandstones referred to the eqalulik formation (figs 82, 83, see below). kangilia formation revised formation history. the kangilia formation was established by rosenkrantz (1970) to encompass the danian (lower paleocene) conglomerate-based, marine mudstone-dominated succession that overlies upper cretaceous mudstones (itilli formation of this paper) with an angular unconformity on northern and central nuussuaq. based largely on the content of fossils, he divided the formation into four members, from base to top: the conglo merate member, a conspicuous coarse-grained, clastsupported conglomerate unit, c. 50 m thick, the fossil wood member, comprising black mudstones with a sandstone unit in the lower part, c. 425 m thick, the thyasira member consisting of sandstones and black mudstones with tuff beds, c. 35 m thick and the propea mussium member made up of black mudstones, locally with intercalated sandstones, c.100 m thick. without providing details, rosenkrantz (in henderson 1969 fig. 4) included the entire 300 m thick mudstonedominated succession at ataata kuua (near atâ) in the kangilia formation. furthermore, the distribution of the kangilia formation was shown in a schematic section through the nuussuaq basin that was presented by henderson et al. (1976 fig. 303). the two upper members of the formation were also reported from alianaat sunnguaq, tupaasat and ataa on the south coast of nuussuaq (rosenkrantz 1970). the kangilia formation including the basal conglomeratic member (annertuneq conglomerate member) is formally defined here. the ‘oyster-ammonite conglo merate’ (birkelund 1965) is formally defined as the oyster–ammonite conglomerate bed. the fossil wood, thyasira and propeamussium members are abandoned. the strata previously assigned to the latter two are now included in the new eqalulik formation (see below). on central nuussuaq, dam et al. (2000) mapped a unit in the kangilia formation that was informally termed the danienrygge member; this unit is not formally recognised here. on itsaku, svartenhuk halvø, j.g. larsen & pulvertaft (2000) suggested that the poorly dated ?upper cam panian/maastrichtian to paleocene succession may be equivalent either to both the itilli and kangilia formations (i.e. campanian to paleocene) or to the kangilia for mation alone (i.e. upper maastrichtian to paleocene), on account of a tentative correlation of two major conglomerate horizons with two discrete tectonic events recognised on nuussuaq. zircon provenance data show that the detrital zircon population of this succession has a distinct and narrow age range peaking at 1870 ma whereas ages of zircons in the underlying sediments of the upernivik næs formation show a greater spread with a distinct peak at 2750 ma (scherstén & sønderholm 2007). since there was only one major change in zircon provenance in the section involving a shift to a unique, single point source at the boundary between the upernivik næs and the overlying deposits, the suggestion is that only one tectonic event is recorded in the section. the entire ?upper campanian/maastrichtian to paleocene succession on itsaku may thus be assigned to the kangilia formation. name. after the headland of kangilia on the north coast of nuussuaq (fig. 74). distribution. the formation is exposed at ataata kuua, ivisaannguit, tupaasat, nuuk killeq and alianaats unnguaq on the south coast of nuussuaq (figs 2, 6, 40; a.k. pedersen et al. 1993), in the northern part of agat dalen on central nuussuaq (fig. 113), in the tunorsuaq valley on western nuussuaq (where it has also been drilled in the gant#1 well) and along the north coast of nuussuaq (fig. 74), and on itsaku on eastern svartenhuk halvø (fig. 73). the formation has been drilled on western nuussuaq in the gro#3 well (fig. 65). type section. the exposures at kangilia on the north coast of nuussuaq are designated as the type section (figs 87, 88). the base of the type section is located at 70°44.90´n, 53°25.33´w. reference sections. well-exposed reference sections occur at annertuneq (fig. 72; nøhr-hansen & dam 1997), 105 ataata kuua (figs 14, 89; dam & nøhr-hansen 2001) and at ivisaannguit west of ataata kuua. the formation was drilled and cored in the gant#1 well in tunorsuaq between 35 m and 256 m (fig. 76) and drilled in the gro#3 well between 728 m and 959 m (fig. 67). thickness. the kangilia formation varies in thickness from 440 m at kangilia (fig. 87) to c. 400 m at ataata kuua (fig. 15), 230 m in the gro#3 well (fig. 67), to just 75 m on central nuussuaq (h.j. hansen 1970). lithology. on northern nuussuaq the kangilia formation comprises a basal approximately 85–140 m thick conglomeratic unit overlain by a mudstone-dominated succession. the conglomeratic unit is here formally established as the annertuneq conglomerate member. in the lower part it is composed of amalgamated conglomerate beds that towards the top become finer-grained and separated by mudstone intervals (figs 72, 76, 87; for detailed description, see below). the succession overlying the basal conglomerate unit consists mainly of dark mudstones with thin beds of fineto coarse-grained sandstone (figs 76, 87, 88). the mudstones are weakly laminated; locally some of the laminae are graded. the sandstone beds have sharp bases and are usually normally graded. the sandstones are structureless or parallel-laminated. ferroan carbonate concretions and fragments of fossil wood are common at various levels. 60 100 120 140 120 180 80 200 220 240 260 280 320 340 360 380 400 420 440 460 480 500 520 540 560 580 600 620 640 660 680 700 720 740 760 780 800 m a.s.l. 300 k an gi lia f or m at io n k an gi lia f or m at io n eq al ul ik f or m at io n va ig at f m it ill i f or m at io n a nn er tu ne q c on gl om er at e m em be r u m iiv ik m em be r clay silt sand pebbl. vf f m c vc f m c clay silt sand pebbl. vf f m c vc f m c ★ ★ ★ ★ k/t t t unconformity po or ly ex po se d su bm ar in e ca ny on po or ly e xp os ed po or ly e xp os ed po or ly e xp os ed po or ly e xp os ed fig. 87. type section of the kangilia formation at kangilia on the north coast of nuussuaq. the log is generalised due to intermittent exposures of the mudstone intervals. top of section is at danienrygge; for location, see fig. 74. k/t, cretaceous– palaeogene boundary. see also figs 88 and 118. for legend, see plate 1. 106 at ataata kuua on the south coast of nuussuaq, a basal sandstone unit with scattered mudstone clasts and transported concretions up to 1.5 m in diameter rests on a deeply eroded surface above lowest campanian and older atane formation (figs 14, 15). the unit forms the base of a 107 m thick fining-upward succession; the sandstone unit grades up into thinly interbedded sandstones and mudstones that are capped by mudstones with scattered basement pebbles interbedded with sandstone lenses and layers (fig. 89; pulvertaft & chalmers 1990; dam & nøhr-hansen 2001). the sandstones are largely structureless, but may grade into paralleland cross-laminated sandstone. convolute bedding is developed locally. the sandstones occur as sheets in composite bedsets in the basal 11 m of the succession, or as single beds that fill wide shallow scours in the middle and upper part (fig. 89). the composite sandstone bedsets are succeeded by approximately 50 m of interbedded sandstones and mudstones consisting of sharply based laminae and beds that grade upwards from coarse-grained to fine-grained sandstone capped by silty mudstone (fig. 89). these lithologies may form sharply based fining-upward units up to 1.5 m thick, restricted to lenticular bodies a few tens of metres wide (fig. 89). erosional discordances are common within the lenticular bodies and internally they are made up of small amalgamated fining-upward successions. the interbedded sandstones and mudstones are succeeded by a 56 m thick unit dominated by massive mudstone containing scattered sand grains and few pebbleand cobble-sized basement clasts (fig. 89). the mudstone beds are interbedded with laminated mudstones with sand streaks (fig. 90). conglomerates and sandstones form an up to 40 m thick unit in the middle of the ataata kuua exposure (figs 14, 15, 89), with sedimentary facies very similar to that seen in the prominent conglomerate unit on the north coast of nuussuaq. this unit is overlain by approximately 250 m of generally poorly exposed mudstone (fig. 14). annertuneq conglomerate member itilli formation umiivik member eqalulik formation kangilia formation fig. 88. outcrop of the itilli, kangilia and eqalulik formations at kangilia. the snowline at c. 800 m corresponds roughly to the base of the volcanic vaigat formation. for location, see fig. 74. 107 just west of ataata kuua, several levels of coarsegrained sandstones and conglomerates are present within the upper mudstone succession (fig. 91). larger clasts include sandstone and mudstone intraclasts, clay ironstone concretions and kaolinised gneiss clasts. dykes of injected sand are also seen. farther west these coarsegrained levels disappear. in agatdalen on central nuussuaq, the formation is generally poorly exposed and dominated by mudstones. a paleocene conglomerate unit about 5 m thick was described as the oyster-ammonite conglomerate by birkelund (1965) and rosenkrantz (1970). it is mainly composed of derived, highly fossiliferous concretions and is here formally established as the oyster–ammonite conglomerate bed. fossils. along the north coast of nuussuaq, fossils are rare in the kanglia formation apart from teredo-bored fossil wood (mathiesen 1961; rosenkrantz 1970). a very 150 m 100 110 120 130 140 50 60 70 80 90 0 10 20 30 40 si pebble cobblesandclay k an gi lia f or m at io n a ta ne fm fig. 89. sedimentological log showing the lower part of the kangilia formation at ataata kuua. slightly modified from dam & nøhrhansen (2001), see also fig. 15. for location of section, see fig. 40; for legend, see plate 1. fig. 90. dark grey mudstones interbedded with thin sandstone beds of the kangilia formation at ataa (for location, see fig. 40). the illustrated section is c. 5 m thick. 108 small number of in situ ammonites (hoploscaphites aff. h. angmartussutensis) and echinoderms have been found in a sandstone bed in the lower part of the formation and in scattered concretions (nøhr-hansen & dam 1997; kennedy et al. 1999). palynomorphs are common and are referred to five zones from the wodehouseia spinata zone to the palaeocystodinium bulliforme zone (nøhrhansen 1997b; nøhr-hansen et al. 2002). the oyster–ammonite conglomerate bed in agatdalen is made up of highly fossiliferous concretions containing a mainly maastrichtian fauna of ammonites although concretions with campanian and danian species are also present. the matrix of the conglomerate contains danian oysters and other bivalves, gastropods, crustaceans and rare specimens of other fossils (birkelund 1965; rosenkrantz 1970). on the south coast of nuussuaq, fossils are very scarce in the kangilia formation and only indeterminate thinshelled bivalves and gastropods have been found in the mudstones at ivisaannguit. the trace fossil planolites isp. and escape burrows occur locally in the sandstone beds. depositional environment. the presence of dinocysts, echinoderms and ammonites indicates a marine depositional environment and the mudstones and sandstones dominating the formation are interpreted to have been deposited from waning, low-density turbidity currents whereas the conglomeratic units were deposited from fig. 91. mudstones interbedded with sandstone and channellised conglomerates (arrows point to base of conglomerate-filled channels). the kangilia formation at ataa; for location, see fig. 40. height of the section is c. 50 m. 109 channellised, high-density currents. some of the mudstones may also have been deposited from suspension. the conglomerates and sandstones at the base of the kangilia formation exhibit a characteristic suite of facies that indicate deposition from catastrophic flows related to high-density, turbidity currents and debris flows. on the north coast of nuussuaq deposition took place in a slope environment, probably in a major submarine canyon setting (figs 76, 87). at ataata kuua, the lower part of the formation was deposited in a submarine canyon (figs 15, 92) whereas the mudstones in the upper part were deposited by dilute turbidity currents in an unconfined slope setting. the basal erosional unconformity represents a transverse section through the submarine canyon (figs 15, 92; dam & nøhr-hansen 2001). the thin lenticular fining-upward successions are interpreted as minor turbidite channel deposits. dinocysts and macrofossils suggest general marine conditions during deposition. along the south coast of nuussuaq, at ataata kuua and ivisaannguit, no marine palynomorphs have been recorded (mcintyre 1993), suggesting a dominant terrestrial input. boundaries. the lower boundary is an erosional unconformity that separates the kangilia formation from the underlying atane formation on southern nuussuaq (fig. 15) and from the itilli formation on western, central and northern nuussuaq (figs 14, 87, 88). along the north coast of nuussuaq, in the gant#1 and gro#3 wells and at ataata kuua, the lower boundary is picked at the base of a thick conglomerate unit that cuts down into the underlying deposits of the itilli and atane formations. at all these localities, the lower boundary marks the base of a major submarine canyon system (fig. 92; dam & nøhr-hansen 2001). where the basal conglomerate is absent, the unconformity between the itilli and atane formations is difficult to locate using lithological criteria (mudstone overlying mudstone) and has to be picked on the basis of biostratigraphic data. on itsaku (svartenhuk halvø), there is an angular unconformity between the kangilia formation and the underlying sediments of the upernivik næs formation. at this locality, a thick conglomerate unit occurs at the base of the formation. elsewhere on svartenhuk halvø, the formation locally onlaps the basement or is in faulted contact with the umiivik member of the itilli formation (fig. 73; j.g. larsen & pulvertaft 2000). the kangilia formation is overlain by the eqalulik formation on northern and western nuussuaq (figs 87, 118), the agatdal formation on central nuussuaq and possibly on itsaku (fig. 113), and by the quikavsak formation on southern nuussuaq (fig. 99a). geological age. the dinocyst flora suggest a late maa strichtian to danian age for the kangilia formation on the north coast of nuussuaq and a danian age in the central part of nuussuaq (fig. 16; j.m. hansen 1970; nøhrhansen 1996, nøhr-hansen & dam 1997; nøhrhansen et al. 2002). the type section of the kangilia formation provides a well-exposed section across the cretaceous–tertiary (k/t) boundary in a clastic, deep-water setting (fig. 87; nøhr-hansen & dam 1997). ammonites in the lowermost part of the formation indicate a late maastrichtian age (kennedy et al. 1999) and coccoliths indicate that most of the section above the k/t boundary has a nanofossil np3–4 zone age and consequently that np2 and b submarine canyon filling a submarine canyon incision ? ? deltaic sandstone turbidite sandstone deltaic mudstone clasts turbidite mud-rich heterolith fig. 92. conceptual model for the formation of the submarine canyon fill of the kangilia formation at ataata kuua. from dam & nøhr-hansen (2001 fig. 4). the canyon formed in response to listric faulting, detachment and collapse of the hanging wall (a). the steeply dipping part of the detachment surface of the footwall acted as the canyon wall and seems not to have been eroded to an appreciable extent. during a subsequent transgression, the canyon was filled with turbidite deposits (b). colours reflect interpreted depositional environments (see plate 1). 110 part of the np1 zone are probably missing or very condensed (nøhr-hansen et al. 2002). the k/t boundary was also cored by the gant#1 well (fig. 76). on central nuussuaq, danian bivalves and gastropods in the matrix of the oyster–ammonite conglomerate member indicate that this part of the formation cannot be older than danian. on the south coast of nuussuaq at ataata kuua, palynomorphs (mainly spores and pollen) suggest a late maastrichtian – early paleocene age for the formation (mcintyre 1993). correlation. the upper, paleocene part of the kangilia formation is coeval with the agatdal and quikavsak formations (fig. 16). subdivision. at the type locality of the kangilia formation, the annertuneq conglomerate member forms the basal unit of the formation (figs 76, 87). in agatdalen, a conglomerate unit containing abundant, derived concretions, the so-called ‘oyster-ammonite conglomerate’ of birkelund (1965), is now formally established as the oyster–ammonite conglomerate bed. annertuneq conglomerate member redescribed member history. this coarse-grained unit was established by rosenkrantz (1970) as the conglomerate member – the basal member of the danian kangilia formation in the type section on the north coast of nuussuaq (figs 72, 93, 94). rosenkrantz (1970) correlated it with his basal danian conglomerate in agatdalen, a unit also referred to as the oyster-ammonite conglomerate by birkelund (1965) and h.j. hansen (1970). based on new data from the north coast of nuussuaq and subsurface data from the gant#1 well, a more thorough and formal description is presented here. the member is retained for the sake of continuity and also because this impressive unit is distinctive as one of the relatively few conglomerates in the nuussuaq basin. name. the name is derived partly from the reference locality close to the type section and partly from its historical name based on the lithology of the unit. distribution. the annertuneq conglomerate member forms a conspicuous unit along the north coast of nuus suaq between niaqorsuaq and kangilia. it has been drilled in the gant#1 well in the tunorsuaq valley (figs 74, 76), and is assumed to be present locally in the subsurface at the base of the kangilia formation. type section. the type section of the annertuneq conglo merate member is the same as for the kangilia formation (figs 74, 87, 88). the base of the member is located at 70°44.54´n, 53°25.11´w. reference sections. a well-exposed section of the lower part of the member is accessible in the western gully leading to annertuneq, a few kilometres west of kangilia (figs 72, 93, 94). the member was cored in the gant#1 well, from 256.0 m to 100.0 m (fig. 76). itilli fm annertuneq conglomerate mb fig. 93. the annertuneq conglomerate member (c. 40 m thick) of the kangilia formation at annertuneq. for location, see fig. 74. 111 thickness. the annertuneq conglomerate member varies in thickness from c. 85 m on the north coast of nuussuaq (fig. 87) to 140 m (excluding igneous intrusions) in the gant#1 well (fig. 76; see also a.k. pedersen et al. 2006b). lithology. at the reference section at annertuneq, the lower part of the conglomerate unit is composed of amalgamated up to 8 m thick lenticular beds containing pebbleto boulder-sized clasts up to 2 m in diameter (fig. 94). clasts include quartz (40%), quartzitic grey sandstones (35%), intraformational mudstones and concretions (10%), quartzitic red sandstones (7%), chert (3%) and others (5%); ‘others’ include precambrian gneiss and granite clasts, and clasts of silicified limestone that may be pisolitic or contain ordovician gastropods. the conglomerates also include coalified wood fragments and poorly preserved marine bivalves and gastropods. individual conglomerate beds can be divided into structureless, structureless to normally graded, or inversely graded, clast-supported conglomerate facies. the conglomerates are poorly sorted and always contain a matrix of medium to very coarsegrained sandstone or sandy mudstone. the sandstones are predominantly structureless, but in some cases the uppermost parts of the beds show well-developed parallel lamination and cross-lamination. dish structures and escape burrows are common. in the type section at kangilia, the lower part of the member consists of a 20 m thick unit of amalgamated conglomerate beds comprising pebbleto boulder-sized clasts in a poorly sorted mediumto very coarse-grained sandstone matrix. the conglomerate units fine upwards and sandstone becomes the dominant lithology towards the top (fig. 87, 95). fossils. scarce, strongly worn bivalves and gastropods and occasional, coalified wood fragments are present in the member. depositional environment. both at the type section and in the gant#1 well, the conglomerates were deposited from debris flows and high density turbidite currents and sandstones from highand low-density turbidity currents. it has previously been suggested that deposition took place in a fluvial environment (h.j. hansen 1970), but the depositional facies and inferred processes, the upward transition into fully marine deposits and the presence of marine dinocysts in the interbedded mudstones (nøhr-hansen 1997b) indicate a marine depositional environment. the considerable increase in the thickness of the member in the gant#1 well compared to the exposures situated at annertuneq and kangilia, c. 6 km to the north, suggests that deposition took place in a major submarine canyon and that the gant#1 well is situated in a more axial position than the north coast exposures. boundaries. the lower boundary of the annertuneq conglomerate member is developed as a major erosional unconformity. on the north coast of nuussuaq,this separates mudstones of the itilli formation below from the conglomerates of the annertuneq conglomerate member above (fig. 72). the upper boundary is gradational at all localities and is placed at the top of the uppermost, very coarse-grained sandstone bed (figs 76, 87, 93, 95). geological age. the annertuneq conglomerate member does not contain fossils of biostratigraphic significance. the age of the member is, however, constrained by its position between marine mudstones. the umiivik mem ber of the itilli formation (below) has a late campanian a nn er tu ne q c on gl om er at e m em be r it ill i f or m at io n k an gi lia f or m at io n fig. 94. close-up of the sharp, erosional base of the annertuneq conglomerate member at annertuneq. person for scale. for location, see fig. 74. 112 age (j.m. hansen 1980b; nøhr-hansen 1996, 1997b; dam et al. 1998c; nøhr-hansen et al. 2002). the mudstones of the remaining kangilia formation (above) are of late maastrichtian – paleocene age. the position of the maastrichtian–paleocene boundary within the kang ilia formation varies between localities, from 210 m above the base of the annertuneq conglomerate member (nøhr-hansen & dam 1997), c. 70 m above the base of the member at kangilia (fig. 87), to a position close to the top of the member in the gant#1 well (nøhrhansen 1997b). this suggests an ?early to late maa strichtian age for the member. oyster–ammonite conglomerate bed new bed history. in agatdalen on central nuussuaq, a conglomerate of danian age containing a huge number of derived concretions has been referred to as the oyster-ammonite conglomerate (e.g. birkelund 1965). h.j. hansen (1970) referred this conglomerate to the thyasira member of rosenkrantz (1970). for historical reasons and in order to distinguish this fossiliferous conglomerate unit from the annertuneq conglomerate member, it is now formally established as a bed and named the oyster –ammo nite conglomerate bed. name. the bed is named after the rich occurrence of reworked oysters and ammonites in the conglomerate. distribution. the bed has only been described from three localities in agatdalen (fig. 113). type locality. the bed is best exposed in the western river bank of the agatdalen river, south-east of sill sø (figs 96, 113). the type locality is located at 70°34.62´n, 53°04.50´w. thickness. the bed is approximately 5 m thick (rosen krantz 1970). annertuneq conglomerate member fig. 95. upper boundary (dashed line) of the annertuneq conglomerate member of the kangilia formation at kangilia showing a gradual transition from conglomerates and sandstones to mudstones. for location, see fig. 74; person for scale. 113 lithology. the oyster–ammonite conglomerate bed is a clast-supported conglomerate dominated by numerous, calcareous concretions and dark mudstone boulders eroded from maastrichtian and campanian deposits (fig. 96; h.j. hansen 1970). the matrix consists of sandy mudstone containing danian oysters and bivalves (see below). in places, calcareous concretions have been formed within the conglomerate (floris 1972). fossils. the fossils in the concretions are reworked and include ammonites, oysters and other bivalves, gastropods, crustaceans and corals (birkelund 1965; rosenkrantz 1970; floris 1972). they show that the concretions were eroded from maastrichtian deposits and represent a rich fauna including hypophylloceras (neophylloceras) groenlandicum birkelund 1965, saghalinites wrighti birkelund 1965, baculites cf. b. meeki elias 1933, scaphites (disco s caphites) waagei birkelund 1965, and s. (d.) angmartussutensis birkelund 1965 (hoploscaphites nowak 1911 according to kennedy et al. 1999). in addition scaphites (hoploscaphites) cf. s. (h.) ravni birkelund 1965, s. (h.) cf. s. (h.) greenlandicus donovan 1953, s. cobbani birke lund 1965 and s. rosenkrantzi birkelund 1965 occur very occasionally and indicate that some concretions were derived from the campanian (kennedy et al. 1999). corals include stephanocyathus sp. and caryophyllia agatdalensis (floris 1972). the mudstone matrix includes oysters and other bivalves that have also been described from the lower part of the overlying eqalulik formation on the north coast of nuussuaq (thyasira member of rosenkrantz (1970)). dinocysts are present in both the reworked concretions and the matrix of the conglomerate. depositional environment. deposition took place from debris flow(s) in a marine environment. boundaries. the lower and upper boundaries of the bed are not well exposed. the lower boundary is inferred to be an erosional unconformity separating the conglomerate from underlying mudstones of the campanian itilli 4 m a see b b fig. 96. a: type locality of the oyster– ammonite conglomerate bed on the western bank of the agatdalen river. for location, see fig. 113. b: clast in the polymodal, matrix-supported conglomerate (paraconglomerate) in the oyster– ammonite conglomerate bed. hammer (encircled) for scale. 114 formation. the conglomerate is overlain by black bituminous mudstones of the kangilia formation (fig. 113). geological age. the oysters and other bivalves in the matrix of the conglomerate indicate a danian age (birke lund 1965). the fauna has species in common with the thyasira member of the kangilia formation of rosen krantz (1970), here defined as the eqalulik formation, and the conglomerate was referred to the lower to middle danian by rosenkrantz (1970). dinocysts from the mudstone matrix also suggest a danian age. kennedy et al. (1999) referred the ammonites to the early maastrichtian. the age of the oyster–ammonite conglomerate bed is constrained by its danian fossils and the early late danian age of the agatdal formation, which overlies the kangilia formation. correlation. the bed is referred to the kangilia formation based on the general danian age and the lack of volcani clastic material (the presence of which characterises the eqalulik formation) in both the conglomerate and the overlying mudstone. quikavsak formation new formation history. the strata comprising the quikavsak formation as defined here were briefly described by steenstrup (1874 p. 79) and later in detail by koch (1959) who included them in his quikavsak member of the tertiary upper atanikerdluk formation. the strata are now established as a separate formation. a major, fluvial channellised sandstone unit that koch (1959 fig. 5) previously included in the top of the atane formation at quikassaap kuua is now tentatively also included in the quikavsak formation. at nassaat, a tributary on the south-eastern side of agatdalen, a poorly exposed conglomerate unit that had been assigned to the sonja member (agatdal formation) by rosenkrantz (1970) and to the quikavsak member by koch (1959) is here re-assigned to the qui kavsak formation. dam (2002) suggested a division of the quikavsak formation into four members reflecting various depositional stages during infill and drowning of the incised valley system. dam’s three lower members – the tupaasat, nuuk qiterleq and paatuutkløften members – are defined formally below. the asuup innartaa member is no longer recognised, and the deposits on nuussuaq are now included in either the eqalulik or the atanikerluk formations whereas the deposits on disko are now referred to the kussinerujuk member of the itilli for mation (see below). name. the member is named after the stream of quikavsaup kûa (now spelled quikassaap kuua) near atanikerluk (fig. 40). the spelling of the name is taken from koch (1959). distribution. the formation is exposed along the south coast of nuussuaq, from nuuk killeq in the west to saqqaqdalen in the east. on central nuussuaq, the formation is exposed at nassaat (fig. 2). paatuutkløften mb nuuk qiterleq mb tupaasat mb paatuutkløften mb nuuk qiterleq mb tupaasat mb atane fm ss q ui ka vs ak f m fig. 97. the type locality of the quikavsak formation, east of nuuk qiterleq on the south coast of nuussuaq (for location, see figs 2, 40). the quikavsak formation overlies the atane formation and is overlain by a very thin eqalulik formation and hyaloclastite breccias of the vaigat formation. a volcanic sill (s) has been intruded between the quikavsak and the eqalulik formations. the thickness of the quikavsak formation in this outcrop is c. 180 m. 115 on svartenhuk halvø, on the east slope of firefjeld, a thin (30 m) paleocene conglomeratic unit is present between the itilli formation and the hyaloclastic rocks of the vaigat formation (fig. 73; j.g. larsen & pulvertaft 2000). this unit may be correlated with either the agatdal formation or the quikavsak formation. type section. originally, koch (1959) chose the exposures at the stream of quikassaap kuua as the type locality for his quikavsak member (fig. 40). at this locality, however, only two of the members of the quikavsak formation are exposed (tupaasat and nuuk qiterleq members) and the formation has an atypical appearance (fig. 100). the best exposures and the most complete development of the formation occur between paatuut and nuuk killeq, south nuussuaq. the type section is located on the coastal slope just east of nuuk killeq below the mountain point 1580 (figs 6, 97, 98, 99a; a.k. pedersen et al. 1993). the type section is located at 70°20.75´n, 53°08.75´w. reference sections. reference sections occur at ivisaannguit, paatuut and quikassaap kuua (figs 40, 99b, 100, 108). thickness. the channellised nature of the deposits forming the quikavsak formation results in highly variable thicknesses. the formation is up to 180 m thick at the type locality. at ivisaannguit it is up to 116 m thick, on the west side of ataata kuua it is up to 135 m, on the east side of ataata kuua below the mountain of ivissussat qaqqaat it is more than 100 m, at paatuut up to 162 m and at quikassaap kuua it is up to 70 m thick. between these sections, the formation is only a few metres thick or is absent altogether. lithology. along the south coast of nuussuaq the formation is divided into three lithological units that are given the rank of members (dam & nøhr-hansen 2001; dam sandclay si pebbles 10 kangilia fm n uu k q ite rl eq m b pm tu pa as at m b q ui ka vs ak f or m at io n a ta ne f or m at io n 0 20 30 40 50 60 70 80 90 100 110 120 m fault fig. 98. type section of the quikavsak formation (and the tupaasat and nuuk qiterleq members) at tupaasat, east of nuuk qiterleq; for location, see figs 2, 6, 97. the upper member of the formation (pm, paatuutkløften member) has not been measured at this locality. the lenticular mud-rich body illustrated on fig. 99a just southeast of the section line is absent in this section, probably due to erosion at the charred base at 39 m. modified from dam & nøhrhansen (2001). note that the position of this section was erroneously stated to be at ivisannguit in dam (2002). for legend, see plate 1. 116 60 0 m .a .s .l. 40 0 20 0 40 0 20 0 20 0 40 0 60 0 80 0 10 00 12 00 14 00 m n w se pa at uu tk lø ft en m b n uu k q ite rl eq m b tu pa as at m b 60 0 m .a .s .l. k an gi lia f or m at io n fi g. 98 pe ak 1 58 0 0 25 0 75 0 m .a .s .l. m .a .s .l. 50 0 75 0 50 0 50 0 75 0 10 00 m a ta ne f or m at io n va ig at f or m at io n q ui ka vs ak f or m at io n a ta ne f or m at io n va ig at f or m at io n q ui ka vs ak f or m at io n n e sw tu pa as at m b n uu k q ite rl eq m b pa at uu tk lø ft en m b pa at uu tk lø ft en tu pa as at t ho le iit ic p ic ri te h ya lo cl as tit es (c ov er ed b y sc re e) pi cr ite in tr us iv es c on te m po ra ry w ith t he v ai ga t fm m ar in e m ud st on es ( eq al ul ik f m ) d el ta ic a nd p ro -d el ta ic m ud st on es an d co al s (a ta ne f m ) be dd in g pl an es t hi n es tu ar in e m ud st on es (p aa tu ut kl øf te n m b) la cu st ri ne m ud st on es an d sa nd st on es ( n uu k q ite rl eq m b) lo w -e ne rg y m ud st on es a nd s an ds to ne s (t up aa sa t m b) d ee p m ar in e m ud st on es ( k an gi lia f m ) d el ta ic s an ds to ne s (a ta ne f m ) es tu ar in e sa nd st on es ( pa at uu tk lø ft en m b) fl uv ia l s an ds to ne s (p aa tu ut kl øf te n m b) fa ul t fl uv ia l c on gl om er at es a nd sa nd st on es ( tu pa aa sa t m b) ba 117 2002). the lower, markedly erosional unit dominated by coarse-grained sandstones succeeded by a middle unit comprising sandand silt-streaked mudstones and sandstones with numerous in situ and drifted remains of trees. the upper unit also has an erosional base and consists of mediumto coarse-grained sandstones arranged in an overall finingand thinning-upward succession, locally with a basal boulder conglomerate bed. for detailed descriptions, see members below. fossils. the sandy parts of the formation are rich in coal fragments and pieces of coalified wood. the fine-grained parts of the formation are rich in clay ironstone with plant fossils, sideritic silty shale with plant fossils, mostly as impressions, and in situ coalified tree trunks. plant fossils belong to the ‘upper atanikerdluk a’ flora of heer (1883a, b); fossil insects and ostracods have also been described (koch 1959). at the top of the section at paatuut, about 5 m below the basalts, a few marine molluscs have been found (koch 1959). early paleocene oysters (ostrea sp.) have been reported immediately west of paatuutkløften and east of nuuk killeq (koch 1959). palynomorphs are common, but only few marine dino cysts are present. depositional environment. the quikavsak formation represents fluvial to estuarine deposition in a series of faultcontrolled, incised valley systems that formed during two phases of uplift of the nuussuaq basin (dam & sønderholm 1998; dam & nøhr-hansen 2001; dam 2002). the mudstone member overlying the basal sandstone member in the type section (fig. 98) is probably related to a quiescent period between the tectonic events resulting in the development of low-energy depositional (lacustrine) areas within the valleys. boundaries. the lower boundary of the formation is a major erosional unconformity formed during valley incision, cutting at least 190 m down into the underlying deposits (figs 99, 109). east of ataata kuua, the unconformity separates turonian – lower campanian deltaic sandstones and mudstones of the atane formation from the coarse-grained deposits of the quikavsak formation. the contact with the underlying atane formation is generally easily recognised with the exception of the locality at quikassaap kuua where the boundary relationships are somewhat enigmatic (see below under tupaasat member). on the west slope of ataata kuua and westwards to ivisaannguit and in the type section, facing page: fig. 99. photogrammetrically measured sections of the incised valley fills of the quikavsak formation. a: longitudinal section of the tupaasat and paatuukløften incised valleys in the area around the type section (see figs 97, 98). b: cross-section of the paatuutkløften incised valley from paatuutkløften (see fig. 105). note that the kangilia formation is bounded by faults, and that the atane formation is also faulted. most of the tupaasat member and the nuuk qiterleq member were removed by erosion prior to deposition of the paatuutkløften member. for location of sections, see figs 2, 40. modified from dam (2002; it should be noted that the figure captions for figs 3b and 4 in this paper are switched). atane formation quikavsak formation20 m fig. 100. reference locality of the quikavsak formation at quikassaap kuua (see fig. 40 for location). this locality was selected by koch (1959) as the type locality of his quikavsak member. the channellised sandstone is now referred to the tupaasat member of the quikavsak formation (for explanation, see text). 118 the quikavsak formation cuts down into mudstones of the upper maastrichtian – lower paleocene kangilia for mation (figs 14, 15, 98, 101). west of paatuut, the formation is overlain by marine mudstones of the eqalulik formation whereas east of paatuut it is succeeded by syn-volcanic lacustrine deposits referred to the atanikerluk formation. geological age. the fauna and flora are not age-diagnostic but a danian age is indicated by the stratigraphic position between the kangilia formation, of danian age on southern nuussuaq, and the overlying eqalulik for mation of danian to selandian age (fig. 16; piasecki et al. 1992; nøhr-hansen et al. 2002. correlation. the quikavsak formation deposits can be correlated to the submarine canyon deposits of the agatdal formation exposed on central nuussuaq (koch 1959; dam & sønderholm 1998) and occurring in the subsurface in the gro#3 well (fig. 67) and the gane#1 well (fig. 119). subdivision. the quikavsak formation is divided into three members: the tupaasat, nuuk qiterleq and paatuutkløften members (fig. 99). tupaasat member new member history. strata now referred to the tupaasat member were first but only briefly described by steenstrup (1874 p. 79). the member was described in detail by koch (1959) and referred to as the lower part of the quikavsak member (upper atanikerdluk formation). that part of the section at quikassaap kuua described by koch (1959 p. 17, fig. 5) as “cross-bedded quartz sandstone of a considerable thickness” and assigned by him to the atane formation, is now referred to the quikavsak formation (see fig. 100). name. the member is named after the headland of tupaasat just west of ataata kuua. distribution. the member is exposed along the south coast of nuussuaq as the fill of an incised valley system extending from nuuk qiterleq in the west to saqqaqdalen in the east where it occurs on the western slope of saqqaqdalen close to the vaigat strait (figs 40, 129; koch 1959; pulvertaft 1989a, b; a.k. pedersen et al. 2007a). the tupaasat member is not present in all outcrops of the quikavsak formation. at some localities it was partly or entirely eroded away prior to deposition of the paatuutkløften member (see figs 99b, 109). mud sand gravel mud sand gravel 0 10 50 40 30 20 60 90 80 70 100 110 140 130 120 m 150 160 190 180 170 200 210 m tu pa as at m em be r n q a ta ne f or m at io n k an gi lia f or m at io n q ui ka vs ak f or m at io n fault fig. 101. generalised sedimentological log from ivisaannguit on the south coast of nuussuaq (for location, see fig. 40). the sandstones and conglomerates of the kangilia formation are similar to those seen at ataa (fig. 91). at this locality, the kangilia formation has a faulted contact with the atane formation and is erosionally truncated by the quikavsak formation. nq, nuuk qiterleq member. for legend, see plate 1. 119 type section. the thickest deposits and the best exposure of the member occur between tupaasat and nuuk qiterleq, in the coastal slope below the mountain point 1580 (figs 6, 98, 99, 102). this exposure is designated as the type section. the type section is located at 70°20.75´n, 53°08.75´w. reference sections. there are several well-exposed sections along the south coast of nuussuaq, e.g. in the coastal cliffs around ivisaanguit (dam 2002 fig. 4) and at quikassaap kuua (figs 40, 100). thickness. the member is 78 m thick in the type section, at least 80 m at ivisaannguit, up to 7 m in paatuutkløften, and c. 20 m at quikassaap kuua (figs 40, 100). lithology. at the type locality, the lower part of the tupaasat member consists of giant-scale trough cross-bedded, parallel-bedded and structureless pebbly to cobbly, very coarse-grained sandstone. in the lower part of the member, each trough set is up to 30 m in thickness (figs 98, 99, 102, 103), but set thickness decreases upwards to less than 0.5 m at the top concomitant with a finingupward trend; the sandstones are, however, distinctly different from the cross-bedded mediumto coarsegrained sandstones of the paatuutkløften member (dam 2002). pebbles are well rounded and consist of quartz (80%), gneiss and granite (10%), metasediments (7%) and chert (3%). within the pebbly sandstones in the type section area, an up to 10 m thick unit of mudstone rich in detrital mica is intercalated with lenses of sandstone. the mudstone is bounded at the top by an erosional unconformity giving rise to the lenticular shape of the shale. the shale is succeeded by cross-bedded pebbly sandstones. around ivisaannguit, the member has a similar development to that in the type section. in the eastern part of these exposures, the member is only up to 15 m thick. towards the west, the thickness of the member increases dramatically to at least 80 m where it crosses an early paleocene fault scar (dam 2002 fig. 4). atane formation tupaasat member fig. 102. giant-scale, cross-bedded conglomerates and sandstones of the tupaasat member (quikavsak formation). encircled person for scale. from the coastal slope between tupaasat and nuuk qiterleq (for location, see fig. 2). 120 in paatuutkløften and at the spur between the coastal escarpment and ippigaarsukkløften, the tupaasat member consists of a 7 m thin unit of pebbly, coarseto very coarse-grained sandstones and a pebble to cobble conglomerate (figs 99b, 104). on the eastern slope of quikassaap kuua, a 20 m thick set of cross-bedded mediumto coarse-grained sandstone is present. it is well consolidated and forms a vertical cliff (fig. 100). this cross-bedded sandstone unit was included in the atane formation by koch (1959) in spite of the stronger similarity to the lower part of the quikavsak member farther west than to the fluvial sandstones of the atane formation. this assignment may have been based on the presence of a small unconformity, possibly associated with a fossil podsol profile that separates the sandstones from the overlying heterolithic deposits, which he referred to as the quikavsak member but are here assigned to the nuuk qiterleq member of the quikavsak formation. fossils. fragments of coal and pieces of coalified wood are common in the member. depositional environment. the conglomerates and sandstones of the tupaasat member form a characteristic suite of very thick sandstone beds deposited by major bedforms related to catastrophic flows. the flows were confined to incised fluvial valleys formed during major tectonic uplift of the basin (dam et al. 1998a; dam 2002). from the tupaasat member facies alone it is not possible to determine whether deposition took place in a subaerial or a submarine environment. the overlying sediments, howfig. 103. giant-scale, cross-bedded conglomerates and sandstones of the tupaasat member. encircled person for scale. arrows indicate dip of major foresets. from the coastal slope above ivisaannguit (for location, see fig. 40). fig. 104. tupaasat member conglomerate in paatuutkløften (for location, see fig. 99b). hammer for scale. 121 ever, have a fluvio-lacustrine origin, which makes a subaerial environment most likely for the tupaasat member. boundaries. the lower boundary is a major erosional unconformity that follows the base of the incised valley systems. it cuts across faults that bring lower paleocene mudstones of the kangilia formation against midcretaceous deltaic deposits of the atane formation (figs 98, 99a). the unconformity reflects major tectonic uplift of the basin (dam 2002). where the nuuk qiterleq member is present, the upper boundary is gradational and non-erosional (figs 97, 98). at other localities, the tupaasat member is erosionally truncated by the paatuutkløften member (fig. 99). geological age. the tupaasat member is referred to the early paleocene (danian), its age being bracketed by data from the kangilia formation below and the nuuk qiterleq member above (dam 2002; nøhr-hansen et al. 2002). nuuk qiterleq member new member history. strata now referred to the nuuk qiterleq member are equivalent to the quikavsak member as described by koch (1959) from the type section at quikassaap kuua. moreover, the nuuk qiterleq member includes the middle, very fossiliferous unit of the quikavsak member described elsewhere in the area by koch (1959). name. the member is named after the headland nuuk qiterleq (fig. 2). distribution. the member is locally exposed along the south coast of nuussuaq as the fill of an incised valley system extending from nuuk qiterleq in the west to atanikerluk in the east (fig. 2). it is not recognised in all outcrops of the quikavsak formation due to erosion preceding deposition of the succeeding paatuutkløften member. type section. the member is well exposed between tupaasat and nuuk qiterleq. the type section is the same as that for the tupaasat member (figs 97–99). the type section is located at 70°20.75´n, 53°08.75´w. reference sections. reference sections occur in paatuut kløften and at quikassaap kuua (figs 99b, 100). thickness. in the type section the member is 23 m thick, at ivisaannguit 3–6 m, at paatuut up to 42 m and at quikassaap kuua up to 50 m. the thickness is mainly dependent on the depth of erosion into the member prior to deposition of the overlying sandstones of the paatuutkløften member (fig. 105). lithology. in the type section and in paatuutkløften, the nuuk qiterleq member is composed of up to 6 m thick coarsening-upward successions, composed of siltand paatuutkløften member nuuk qiterleq member atane formation fig. 105. nuuk qiterleq member erosively overlain by the paatuutkløften member in paatuutkløften (see fig. 99b). helicopter (encircled) for scale. 122 sand-streaked mudstones grading upward into fineto medium-grained sandstones (fig. 106a). incipient ripples as well as wave and current ripples and parallel lamination are widespread. in situ vertical coalified trees and drifted tree trunks and plant fossils are very common (fig. 106b). at ivisaannguit, at paatuut and at ippigaar sukkløften, the member is dominated by mudstones with thin sand-streaks. a single, sharply based sandstone bed occurs at ivisaannguit. at quikassaap kuua, the member consists of interbedded siltstones and graded, occasionally parallellaminated, very fine-grained to very coarse-grained sandstones (fig. 107). this interbedded facies alternates with lenticular beds with erosional bases consisting of normally graded, mediumto very coarse-grained trough, crossbedded sandstones arranged in thinning-upward successions up to c. 50 cm thick. the interbedded siltstones and sandstones are rich in plant fossils, clay ironstone concretions, in situ vertical coalified tree trunks, drifted coalified tree trunks and finely disintegrated coal fragments. fossils. the nuuk qiterleq member is very rich in plant fossils and contains most of the species known from the quikavsak formation. the fossil flora from this member and from the lower part of the overlying naujât member (atanikerluk formation) at atanikerluk and a few additional localities is known as the ‘upper atani kerdluk a’ flora of heer (1883a) who recognised 282 species, some of which have been revised by brown (1939), seward (1939), chaney (1951) and koch (1963). koch described the lower paleocene flora in the agatdalen area where he recognised 27 species, of which 19 species were recognised by him as also occurring in the quikavsak formation. depositional environment. sediments of the nuuk qiterleq member were deposited in various environments within an incised valley system. the stacked coarsening-upward cycles in the type section and in paatuutkløften were formed by repeated progradation of shoreface or bayhead deltas into a lacustrine environment (dam 2002). the deposits at quikassaap kuua include lenticular fining-upward successions of sandstones deposited in minor fluvial channels, interbedded with siltstones and graded, very fine-grained to very coarse-grained gravelly sandstones deposited in interchannel areas. the thin graded sandstone sheets were probably deposited during periods of intense overbank flooding. boundaries. the lower boundary is placed at the lithological change from pebbly coarse-grained sandstones of the tupaasat member to the fine-grained deposits of nuuk qiterleq, rich in plant fossils (fig. 99). at most localities, the lower boundary is non-erosional; at qui kassaap kuua, however, the nuuk qiterleq member overlies an erosional surface (koch 1959). paatuutkløften member n uu k q ite rl eq m em be r a b fig. 106. a: thin coarsening-upward successions in the nuuk qiterleq member composed of siltand sand-streaked mudstones passing upwards into fineto medium-grained sandstones. person for scale. from the type locality near nuuk qiterleq (for location, see fig. 2). arrows indicate the location of in situ coalified tree trunks. note the upper, sharp, erosional boundary with the paatuutkløften member. b: in situ vertical, coalified tree trunk from the type section of the nuuk qiterleq member. pencil for scale (arrowed). 123 at most localities, the upper boundary is with the paatuutkløften member (figs 99, 106). at quikassaap kuua, the member is overlain by dislocated (landslipped) mudstones of the atanikerluk formation (koch 1959). geological age. all the plants are interpreted as danian in age (koch 1963). paatuutkløften member new member history. strata now referred to the paatuutkløften member were first described and figured as the quikavsak member by koch (1959) along the south coast of nuussuaq in the paatuut area (koch 1959 figs 20–25), the ataata kuua area (koch 1959 figs 26–30) and at the localities between tupaasat and nuuk qiterleq and on central nuussuaq at nassaat (koch 1959 fig. 37). the strata were described in detail by (dam et al. 1998a; dam & sønderholm 1998; dam 2002). name. the member is named after the paatuutkløften ravine (fig. 40). distribution. the paatuutkløften member is exposed along the south coast of nuussuaq as the fill of an incised valley system, exposed from nuuk qiterleq in the west to ippigaarsukkløften in the east and also at nassaat on central nuussuaq (figs 2, 40). type section. the eastern slope of paatuutkløften (figs 99b, 105, 108). the base of the type section is located at 70°15.27´n, 52°40.65´w. reference sections. reference sections occur on the western slope of ippigaarsukkløften (figs 108, 109), at ivissussat qaqqaat on the east side of ataata kuua (fig. 110), in the south-west-facing cliffs above the west side of ataata kuua (fig. 14) and between tupaasat and nuuk qiterleq below the mountain point 1580 (figs 6, 97, 98, 99; a.k. pedersen et al. 1993). thickness. the member is 145 m thick in the type section. thicknesses in the reference sections are 158 m at ippigaarsukkløften, more than 110 m at ivissussat qaqqaat, 110 m in the coastal cliffs above ataata kuua, up to 25 m at the coastal section above ivisaannguit and 78 m in the section between tupaasat and nuuk qiterleq. the exposures at nassaat are too poor to estimate the thickness of the member. lithology. at ivissussat qaqqaat and ataata kuua the member is initiated by conglomerates composed of gneiss pebbles and boulders (fig. 110). this is followed by monotonous successions of light grey, well-sorted, dominantly planar cross-bedded, mediumto very coarsegrained pebbly sandstones that are present at all localities (fig. 111). these are distinctly different to the succession characterising the tupaasat member (dam 2002). there is a general overall upward decrease in grain size, clast size and set thickness; thin mudstone beds are occasionally present in the uppermost part of the member. coal clasts, finely disseminated coal fragments and sideritic mudstone clasts are common. penecontemporaneous deformation structures are widespread in some beds. along the south coast of nuussuaq between nuuk qiterleq and ippigaarsukkløften, the sandstones become 1 m fig. 107. interbedded siltstones and sandstones of the nuuk qiterleq member on the western slope of quikassaap kuua (for location, see fig. 40). this section was previously the type section of the quikavsak member of koch (1959). 124 0 cl si sand pebbles 10 20 30 40 50 60 70 80 90 100 110 120 130 140 150 m cl si sand pebbles base of incised valley 0 10 20 30 40 50 60 70 80 90 100 110 120 130 140 150 m terminal lobe paatuutkløften ippigaarsukkløften hydrothermal fra cture atane fm eqalulik fm pa at uu tk lø ft en m em be r q ui ka vs ak f or m at io n fig. 108. type and reference sections of the paatuutkløften member at paatuutkløften and ippigaarsukkløften, respectively (for location, see fig. 40; for legend, see plate 1). the dashed lines indicate correlative surfaces between the two sections. 125 light orange in the upper part of the member and crossbeds show double mud drapes and reactivation surfaces draped by mudstone clasts. locally the cross-bedding is bidirectional. the sandstones frequently show convolute bedding; coal fragments, sideritic mudstone clasts and concretions are common. thin mudstone beds with plant debris and heavily bioturbated sandstones occur in the upper part of the succession. fossils. few recognisable plant fossils are present in the mudstones of the member. on the spur between the western slope of the ippigaarsukkløften ravine and the coastal slope, about 5 m below the basalts, a few danian marine molluscs have been found (koch 1959). danian oysters (ostrea sp.) have been found immediately west of paatuutkløften and east of nuuk killeq (fig. 2; koch 1959). these oysters were most probably derived from sandstones in the uppermost part of this member. trace fossils are common to abundant at most localities in the upper part of the member. they include ophiomorpha nodosa, ophiomorpha isp., diplocraterion parallelum and thalassinoides isp. (fig. 108). depositional environment. the sediments of the paatuut kløften member were deposited in an incised valley system formed after an episode of renewed early paleocene faulting. the valley system seems to follow the system generated prior to the deposition of the tupaasat member, resulting in considerable or in some cases complete erosion of the earlier deposits within the incised valleys. deposition of the lower part of the member took place in a fluvial environment during a period of rapidly rising sea level, high river discharge and high sedimentation rates (koch 1959; dam & sønderholm 1998; dam 2002). the upper part of the member is characterised by a change in colour, by sedimentary structures typical of tidal activity and by intense bioturbation, indicating a change to a tidal estuarine environment during infilling of the valleys (dam & sønderholm 1998). atane formation quikavsak formation vaigat formation sill fig. 109. paatuutkløften member as exposed in the ippigaarsukkløften section (for location, see fig. 40). the sandstones of the paatuutkløften member fill a valley incised into deltaic deposits of the atane formation (see fig. 108). thickness of valley fill is c. 180 m; the white arrow indicates the base of the reference section shown in fig. 108. 126 boundaries. an erosional unconformity separates the sandstones of the paatuutkløften member from the underlying incised valley deposits of the nuuk qiterleq or tupaasat members (figs 97, 99, 105). in some cases, the paatuutkløften member is incised directly into the kangilia or atane formations suggesting either that the previous deposits within the valley system were completely eroded away or that new valleys were eroded locally (figs 14, 99, 109, 110). the sandstones of the paatuutkløften member are abruptly overlain by mudstones of the eqalulik formation (figs 108, 112). geological age. the age of the paatuutkløften member is bracketed by the danian age of the underlying kangilia formation and the latest danian to selandian age of the overlying eqalulik formation (see p. 137), suggesting that the age of paatuutkløften member is late np3–np4 (nøhr-hansen et al. 2002). correlation. the paatuutkløften member has been correlated with marine deposits of the agatdal formation on central nuussuaq (koch 1959; dam & sønderholm 1998), and with the agatdal formation in the gane#1 and gro#3 wells (nøhr-hansen et al. 2002). atane formation paatuutkløften member fig. 110. basal gneiss-clast conglomerate of the paatuutkløften member at ivissussat qaqqaat (east slope of ataata kuua; for location, see fig. 40). person for scale. fig. 111. uniformly cross-bedded fluvial sandstones of the paatuutkløften member at ippigaarsukkløften (for location, see fig. 40). the apparent discontinuity in the middle of a section is due to hydrothermal alteration (see fig. 108). encircled person for scale. 127 agatdal formation revised formation history. the agatdal formation was established by rosenkrantz (in: koch 1959 pp. 75–78) to encompass upper danian marine mudstones, sandstones and conglomerates found in the agatdalen area on central nuus suaq; these sediments were further described by h.j. hansen (1970), rosenkrantz (1970) and henderson et al. (1976). in some papers, the formation is referred to as the agatdalen formation. the agatdalen area was the key study area for the nûgssuaq expeditions led by a. rosenkrantz because of the rich fossil content of the sediments (see ‘previous work’). focus was on the fossil assemblages, and the various outcrops (sections) were only shown schematically and described provisionally. the component lithological units of the section were assigned to individual members resulting in a complex lithostratigraphic terminology. four members were recognised within the small area of outcrop: the turritellakløft, andreas, sonja and abraham members. schematic sections from the localities in the agatdalen area are shown in rosenkrantz (1970 figs 4, 5). the outcrops in the agatdalen area are discontinuous, and because many of the sandstone and conglomerate units are channellised – and lateral facies changes are therefore pronounced – the members are difficult to correlate and map in detail (figs 113, 114). a possible correlation between the outcrops of the agatdal formation is shown in plate 3. furthermore, the members cannot be recognised outside the type area and therefore the three lower members are no longer treated as formal units. they are, however, treated as informal members below in order to facilitate correlation between the new data presented here and the older literature and fossil records (see plate 3). the agatdal formation as defined here is entirely prevolcanic. thus, the abraham member, which is synvolcanic, is now included in the new eqalulik formation (see below). rosenkrantz (1970) referred other synvolcanic marine sediments found on nuussuaq to the tuffaceous thyasira and propeamussium members (abandoned here) of the kangilia formation. in agatdalen, the turritellakløft member, from which tuffs are not reported, overlies the propeamussium member (rosenkrantz 1970 fig. 5). the correlation between the preand synvolcanic members of the kangilia and agatdal formations of rosenkrantz is not well explained. strata assigned to the agatdal formation (sonja member) at nassaat north-east of agatdalen (fig. 2) by rosenkrantz (1970 fig. 4c, d) are here assigned to the quikavsak formation, following koch (1959). the tuffaceous deposits of the abraham member are coeval with the volcanic asuk member of the vaigat formation and are therefore now included in the eqalulik formation (see p. 137). a sandstone unit underlying the pillow breccias at kangilia on the north coast of nuussuaq was tentatively assigned to the agatdal formation by h.j. hansen (1970), rosenkrantz (1970) and dam & sønderholm (1998); this unit is regarded here as forming part of the eqalulik formation. eq al ul ik fo rm at io n pa at uu tk lø ft en m em be r fig. 112. the sharp boundary between estuarine sandstones of the paatuutkløften member and the mudstones of the eqalulik formation above. hammer (left) for scale. from ippigaarsukkløften (for location, see fig. 40). 128 subsurface strata, here assigned to the agatdal for mation, were first cored in the gane#1 well drilled by grønarctic energy inc. at eqalulik on western nuussuaq in 1995 (fig. 65). a thick succession of marine paleocene sandstones and conglomerates was drilled in 1996 by the gro#3 exploration well south-east of the kuussuaq river delta (figs 9, 65, 67). this succession was referred to the quikavsak formation by christiansen et al. (1999), but is here referred to the agatdal formation. name. the formation is named after the agatdalen valley on central nuussuaq. distribution. outcrops of the agatdal formation are only known from the agatdalen area (fig. 113; dam et al. 2000). from subsurface data (gro#3, gane#1 and gane#1a wells), the formation is known also to be present on western nuussuaq, west of the kuugannguaq– qunnilik fault. teltbæk a gatda len i i i i i i i i i i i i ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ ■■ 600 600 800 1000 600 800 1000 40 0 60 0 80 0 10 00 400703 oa oa 400704 400702 400701 800 10 00 53°10' 53°05' 53°00' 70°35' 53°10' teltbæk fault agatdalen fault baculiteskløft søndre and nordre jee pklø ft 17 14 12 14 26 scaphitesnæsen turrite llakløft q aa rs ut jæ ge rd al pyram idedal 1 km br aya kløft agatkløft ’sonja lens’ sill sø strike and dip of strata contour in metres vaigat formation eqalulik formation quaternary cover agatdal formation kangilia formation itilli formation, aaffarsuaq member atane formation intrusion ice sea/lake type section (oa) 500 14 oa exposure fig. 113. geological map of the agatdalen area. the positions of the four shallow wells drilled by ggu in 1992 (ggu 400701–400704) are shown. exposures of oyster–ammonite conglomerate bed (oa) shown with dots. from dam et al. (2000) based on rosenkrantz et al. (1974). turritellakløft is also named turritelladal on some maps. for location, see fig. 2. 129 the turritellakløft member and the abraham member of rosenkrantz (1970) were described from the western end of the turritellakløft gorge and qaarsutjægerdal (figs 113–115, plate 3; h.j. hansen 1970), whereas the andreas member was only recognised in turritellakløft (figs 113, 114, plate 3). the sonja member of rosen krantz (1970) is only known from a section in agatkløft, 2 km east of turritellakløft (figs 116, 117, plate 3). on svartenhuk halvø, on the east slope of firefjeld, a thin (30 m) paleocene conglomeratic unit is present between the itilli formation and the hyaloclastic rocks of the vaigat formation (fig. 73; j.g. larsen & pulvertaft 2000). this unit may be correlated with either the agatdal formation or the quikavsak formation. type section. the type section of the agatdal formation is at the so-called ‘store profil’ (or ‘big section’) in turri tellakløft in the north-western end of agatdalen (fig. 114, plate 3, section 1; rosenkrantz 1970 fig. 5; hender son et al. 1976 fig. 312). the type section is located at 70°35.01´n, 53°08.02´w. reference sections. reference sections occur on the southern slopes of agatkløft and on the eastern slope of qaar sutjægerdal (figs 115, 116, 117, plate 3). in the subsurface, a complete section through the forma tion was encountered in the gro#3 well from 718 m to 423 m (fig. 67). no cores were taken but a complete suite of well logs is available (kristensen & dam 1997). the upper part of the agatdal formation was drilled and fully cored in the gane#1a well from 706.7 m to 601.5 m (fig. 119). thickness. the formation varies considerably in thickness from 65 m in the type section to 18 m in qaarsutjægerdal (plate 3). in the gro#3 well, the formation is approximately 250 m thick (excluding igneous intrusions; fig. 67). a br ah am m em be r ‘a nd re as m em be r’ ‘t ur ri te lla kl øf t m em be r’ a ga td al f or m at io n eq al ul ik f or m at io n fig. 114. type locality of the agatdal formation in turritellakløft (rosenkrantz 1970; for location, see fig. 113). the section is approximately 80 m high. in the present paper, the agatdal formation is not subdivided, but the formally abandoned ‘turritellakløft member’ and ‘andreas member’ are shown to provide a link to rosenkrantz (1970). the agatdal formation is overlain by a thin development of the abraham member which is part of the eqalulik formation. for measured section, see plate 3 (section 1). 130 lithology. in the agatdalen area, the agatdal formation comprises mudstones with several thick, lenticular conglomeratic and sandstone units (figs 114–117, plate 3) that rosenkrantz (1970) had already recognised to be partly laterally equivalent. the agatdal formation, as exposed in the junction between agatkløft and qaarsutjægerdal, is characterised by a major, fining-upward succession composed of a basal conglomerate unit grading upwards into alternating arkosic sandstones, conglomerates and mudstones with marine fossils and plant remains (plate 3). the basal conglomerate unit can be followed laterally for about 1 km, but disappears rapidly towards the south-east along the agatdalen river (figs 113, 116, 117, plate 3). the conglomerates also seem to wedge out towards the northwest in agatkløft. the clasts are pebble to boulder grade and are composed of gneiss. the conglomerates form amalgamated beds up to 7 m thick, or alternate with clast-rich, coarse-grained sandstones. the sandstone beds of the overlying succession usually show normal grading and both paralleland cross-lamination. sandstone streaks are common in the mudstones interbedded with the sandstone beds. in the turritellakløft section, the kangilia formation is unconformably overlain by dark grey to black, sandstreaked mudstones with sandstone lenses and sheets which form the base of the agatdal formation (fig. 114, plate 3; rosenkrantz 1970). the lower part of the formation here consists of thinly interbedded mudstones, siltstones and sandstones. the mudstones show graded, parallel lamination. the sandstone beds are paralleland cross-laminated. trace fossils are common (j.m. hansen 1980b). upwards, the sandstone beds become thicker and occur as sheets and lenses with erosional bases. the beds are parallel laminated or show normal grading from coarse-grained to fine-grained sand. locally the sandstones are highly fossiliferous (predominantly turritella sp.). intraformational mudstone and sandstone clasts are common. loose slabs show a preferred orientation of turritella shells; this parallel orientation indicates that the shells were current-worked and may serve as palaeocurrent indicators. in the upper part of the turritellakløft section is a channellised unit of well-sorted, very coarse-grained, tabular to wedge-shaped, cross-bedded sandstone (fig. 114). the top of the agatdal formation is a 1.5 m thick gravelly, very coarse-grained sandstone bed. body and trace fossils are very rare in these sandstones. in qaarsutjægerdal, a matrix-supported conglomerate bed occurs at the base of the formation (fig. 115, plate 3). it is succeeded by well-sorted, fineto mediumgrained sandstones, with few mudstone interbeds. locally the sandstones show paralleland low-angle cross-lamination. bioturbation is intense at the top of the section where ophiomorpha isp. and planolites isp. are common. the sediments are highly fossiliferous; turritella sp. is particularly numerous and the shells show a preferred parallel orientation. in the gane#1 well, the agatdal formation consists of a thick, fining-upward succession composed of a lower sandstone unit grading upwards into a heterolithic unit. eqalulik formation agatdal formation pillow breccia kangilia formation fig. 115. thin development of the agatdal formation in qaarsutjægerdal (for location, see fig. 113). sandstone beds of the agatdal formation overlie mudstones of the kangilia formation with marked angular unconformity. the agatdal formation is overlain by the abraham member of the eqalulik formation. the exposed section below the pillow breccias is 21 m thick. for measured section, see plate 3 (section 4). 131 the cored succession in the gane#1 and gane#1a wells consists of amalgamated, thickly bedded, coarseto very coarse-grained sandstones with normally graded beds in places alternating with thinly interbedded sandstones and mudstones (fig. 119). the upper heterolithic part of the formation consists of mudstones, thinly interbedded sandstones and mudstones, amalgamated sandstones grading upwards into thinly interbedded sandstones and mudstones and muddy sandstones. probable scape burrows are observed locally (dam 1996b). fossils. the agatdal formation is the most fossiliferous of the marine units in west greenland, with more than 500 described macrofossil species (see also ‘previous work’ above). marine fossils in the mudstones (the former turri tellakløft member) include gastropods, bivalves, scleractinian corals, fish, ostracods, coccoliths and foraminifera (bendix-almgreen 1969; rosenkrantz 1970; szczechura 1971; j.m. hansen 1976; kollmann & peel 1983; petersen & vedelsby 2000). dinocysts are rare (j.m. hansen 1980b). the sandstones of the former andreas member are very poor in fossils in the turritellakløft section, whereas certain levels in qaarsutjægerdal are very fossiliferous. turritella sp. is particularly common, but other gastropods together with bivalves, scleractinian corals and spines of echinoids have also been found in the member (rosenkrantz 1970; floris 1972; kollmann & peel 1983; petersen & vedelsby 2000). wood and fragments of leaves and fruits have been collected in this member in qaarsutjægerdal (koch 1963, 1972a, b). the sandstones and conglomerates of the former sonja member are generally poor in marine fossils but locally contain a very rich marine fauna, known mainly from the ‘sonja lens’. the lens was originally 7 m long and 0.7 m thick but was completely excavated by rosenkrantz and his co-workers. the fauna from this bed is dominated by bivalves and gastropods, but also includes scleractinian corals, octocorals, asteroids, crinoids, echinoids, serpulids, brachiopods, bryozoans, scaphopods, crustaceans, fish, foraminifera, coccoliths and palynomorphs (bendixalmgreen 1969; h.j. hansen 1970; rosenkrantz 1970; szczechura 1971; floris 1972; perch-nielsen 1973; henderson et al. 1976; j.m. hansen 1980b; kollmann & peel 1983; collins & wienberg rasmussen 1992; petersen & vedelsby 2000). moreover, the agatdal formation contains a well-preserved macroflora (koch 1963). on western nuussuaq, however, no determinable macrofossils have been found in the agatdal formation. dinocysts are common in samples from the gro#3 and gane#1 wells (nøhr-hansen 1997b, c; nøhrhansen et al. 2002). the trace fossils ophiomorpha isp. and planolites isp. are locally common (see above). depositional environment. in the agatdalen area, the macrofossils are indicative of a relatively shallow-water, marine depositional environment, such as a lower shoreface or inner shelf environment (petersen & vedelsby 2000). in contrast, the sedimentary facies are interpreted to reflect deposition in a deep-water slope environment. fig. 116. basal conglomerate unit of the agatdal formation south of the junction between agatdalen and agatkløft (for location, see fig. 113). note how the conglomerate pinches out towards the left (south). section is approximately 13 m high. for measured section, see plate 3 (section 3). 132 the mudstone-dominated units were deposited from dilute turbidity currents in an unconfined slope setting where the sandstone interbeds possibly represent splay deposits from nearby submarine channels. the thick lenticular, conglomeratic sandstones were deposited by highand low-density turbidite currents in submarine slope channels. the fossils found in these units must therefore have been transported, but probably only over short distances, as they are not worn or fragmented. the presence of basement boulders in the conglomerate close to the base of the formation implies fluvial transport from areas east of the ikorfat fault system (a.k. pedersen et al. 1996 fig. 6). the conglomerates were thus probably deposited in a submarine canyon in front of a major delta at the mouth of the coeval incised fluvial valley system represented by the quikavsak formation. palynological and geochemical investigations of the gro#3 and gane#1 wells also indicate a marine depositional environment (bojesen-koefoed et al. 1997; nøhrhansen 1997c) and the sediments record deposition dominated by high-density turbidity currents. the thickness of the formation in the gro#3 and gane#1 wells and the overall geological setting suggests that deposition took place in a major submarine canyon system on a fault-controlled slope (dam 1996b; kristensen & dam 1997). the small, fining-upward cycles in the upper, heterolithic part of the canyon fill were deposited in small turbidite channels. boundaries. where the lower boundary is exposed in the agatdalen area the agatdal formation rests unconformably on the kangilia formation (figs 113, 115, plate 3). the agatdal formation is overlain by the eqalulik formation (figs 113–115, 119, plate 3) or locally by volcaniclastic breccias of the vaigat formation. in the gro#3 well the lower boundary is probably an erosional unconformity separating coarseto very coarse-grained sandstones of the agatdal formation from fig. 117. exposure of the agatdal formation in agatkløft (for location, see fig. 113; for measured section, see plate 3 (section 2)). the section is approximately 50 m high and, together with the basal conglomerate unit shown in fig. 116, was referred to the ‘sonja member’ by rosenkrantz (1970). the completely excavated sonja lens was part of this section. according to h.j. hansen (1970) the succession is capped by 6 m of concretionary mudstone (covered by scree during the authors’ visit in 1992) underlying volcanic pillow breccias and a sill. 133 heterolithic deposits of the kangilia formation below (fig. 67). the upper boundary is placed at the base of the reworked volcaniclastic sediments or tuffs of the eqalulik formation (fig. 67). age. most fossil groups indicate a paleocene (late danian) age for the formation (fig. 16; bendix-almgreen 1969; h.j. hansen 1970; rosenkrantz 1970; szczechura 1971; kollmann & peel 1983; collins & wienberg rasmussen 1992). the coccolith assemblage indicates a chiasmolithus danicus zone age (np 3 zone; perch-nielsen 1973). the co-occurrence of the foraminifera globoconusa daubjergensis and globigerina compressa indicates a p1c plankton zone age for the agatdal formation (h.j. hansen 1970; berggren et al. 1995; olsson et al. 1999) which corresponds to a late np3 – early np4 age. in the gro#3 well, dinocysts and nannofossils indicate an age not younger than nannoplankton zone np4 (late danian or early selandian; nøhr-hansen 1997c; nøhr-hansen et al. 2002). correlation. the age of the agatdal formation suggests that it is coeval with the major incised valley systems of the quikavsak formation on southern nuussuaq and with most of the upper part of the kangilia formation on the north coast of nuussuaq (figs 11, 15, 16; nøhrhansen et al. 2002). the basal unconformity described between the agatdal and kangilia formations in the agatdalen area may be due to either condensation or nondeposition of nannoplankton zones np1 and np2 in the type section of the kangilia formation. eqalulik formation new formation history. deposits now assigned to the eqalulik formation embrace the marine synvolcanic strata below the basalts of the west greenland basalt group (hald & pedersen 1975) that contain volcaniclastic sandstones and tuffs. on the north coast of nuussuaq and on central nuussuaq, these strata have previously been assigned to the thyasira and propeamussium members of the kangilia formation (rosenkrantz 1970) and the abraham member of the agatdal formation (rosenkrantz 1970). however, the thyasira and propeamussium members are considered inappropriate as they were established solely on the basis of the faunal content. the use of these two members as defined by rosenkrantz (1970) is now abandoned and the strata included in the new eqalulik formation. on the south coast of nuussuaq, a thin unit of marine black shale including bioturbated sandstone beds below the volcaniclastic breccias of the vaigat formation was referred to the asuup inatartaa member of the quikavsak formation by dam (2002) but is here included in the eqalulik formation. strata previously mapped as unnamed paleocene (tertiary) on the north coast of nuussuaq east of the ikorfat fault, in the region around the tunorsuaq valley, and in the itilli and aaffarsuaq valleys are now assigned to the eqalulik formation. name. after eqalulik, a lake close to the location of the gane#1 drill site on western nuussuaq (fig. 65). distribution. the formation is widely distributed below the hyaloclastites of the anaanaa and naujánguit members of the vaigat formation (west greenland basalt group) and, although it is generally poorly exposed, it has been recognised at many localities on nuussuaq (piasecki et al. 1992), locally on the north coast of disko, and locally on svartenhuk halvø (fig. 73; j.g. larsen & pulvertaft 2000). on the north coast of nuussuaq, the formation has been recognised both east and west of the ikorfat fault. it is present in the tunorsuaq valley and has been locally observed in the itilli and aaffarsuaq valleys (floris 1972). west of ataa on the south coast of nuussuaq, the formation forms a thin unit dominated by marine mudstones below the hyaloclastites of the vaigat formation. east of ataa, it is sandwiched between the sandstones of the quikavsak formation and overlying non-marine mudstones referred to the naujât member of the atanikerluk formation (see also koch 1959 pp. 78–79). type section. the outcrop section at kangilia where the eqalulik formation conformably overlies the kangilia formation is very poorly exposed (figs 74, 87, 118). a complete, well-preserved section through the formation was cored from 598 m to 496.5 m in the gane#1 well (figs 65, 119) and this is therefore chosen as the type section (fig. 119). the cores are stored at geus in copen hagen. the gane#1 well is located at 70°28.25´n, 54°00.40´w. reference section. reference sections occur on central nuussuaq in agatdalen (plate 3) and aaffarsuaq (fig 83), and on the south coast of nuussuaq at ippigaar suk kløften (fig. 120). in the gank#1 well, a reference section was cored from 333 m to 114.9 m (fig. 121), and the formation was also cored in the sidetrack well gank#1a from 332.9 m to 218.6 m. the cores are stored at geus. 134 thickness. on the north coast of nuussuaq, the formation ranges from 100 to 160 m thick, in agatdalen it is approximately 12 m thick (h.j. hansen 1970), and in aaffarsuaq it is less than 20 m in thickness. in the gane#1a well, the formation is 102 m thick and it is 218 m thick in the gank#1 well. the very variable thickness of the formation is the result of varying proportions of tuffaceous material (the presence of which defines the formation) and intensive loading and deformation of the mudstones by prograding thick lobes of submarine basaltic breccia beds. lithology. the formation comprises mudstones, tuff beds and volcaniclastic sandstones; the latter lithologies are essential to the definition of the formation (fig. 122). internal structures in the mudstones cannot be recognised due to poor exposure. the thicker sandstone units consist of amalgamated graded beds composed of fineto very coarse-grained, volcaniclastic sandstones with scattered pebble-sized volcanic and mudstone clasts. parallel lamination, current cross-lamination and dish structures have locally been observed within the graded beds. in some areas volcanic material may constitute up to 25% of the section (a.k. pedersen 1978) and fossils are common (rosenkrantz 1970; floris 1972). the tuff beds are up to 7 m thick at kangilia, and tuff beds up to several metres thick are also seen in danienkløft (tunorsuaq) and in koralkløft (ilugissoq; figs 2, 74). most of these tuff beds contain scleractinian corals; these are without preferred orientation, but only slightly worn (floris 1972). in the gane#1 well the formation consists of mudstones, interbedded quartz-rich, siliciclastic and volcaniclastic sandstones and mudstones, muddy sandstones and chaotic beds (dam 1996b). volcanic clasts and rounded basement pebbles and angular mudstone ripup clasts are common, and concretions and plant debris are locally present (fig. 119). several thinningand fining-upward successions have been recognised in the core. these are sharply based and consist of amalgamated normally graded, very coarseto coarse-grained sandstone beds passing upwards into thinly interbedded sandstones and mudstones. occasionally the sandstone beds are internally stratified, showing parallel-lamination or crosslamination, dish structures and soft sediment folds. sandstones are occasionally burrowed and escape burrows eqalulik formation kangilia formation fig. 118. exposure of the eqalulik formation overlying the kangilia formation at kangilia. the photo shows the lower c. 20 m of the formation; for location, see fig. 74. 135 vf f m c 640 600 550 500 m vc sandsi pbclay vf f m c vc sandsi pbclay hyaloclastites 700 650 600 550 m first appearance of hyaloclastite clasts gane#1 va ig at f m a na an aa m b eq al ul ik f or m at io n a ga td al f or m at io n gane#1a fig. 119. type section of the eqalulik formation in the gane#1 and 1a wells on western nuussuaq. for location, see fig. 65; for legend, see plate 1. 136 are present in a few beds. mudstone laminae show normal grading. the chaotic beds in the gane#1 well consist of homogenised mudstones with evenly scattered sand grains, granules, volcanic clasts and concretions. the beds are interbedded with structureless muddy sandstones and slumped mudstones showing contorted bedding (fig. 119; dam 1996b). fossils. the formation has yielded a very rich fauna, particularly in the kangilia section where the concretions in the mudstones include gastropods, bivalves, corals, echinoderms and nautiloids, and also foraminifera, crinoids and serpulids. the bivalve thyasira (conchocele) aff. t. conradi (rosenkrantz) and the pectinid propea mussium ignoratum ravn are especially common in the concretions together with a large thick-shelled echinocorys, hercoglossa groenlandica (rosenkrantz) and stems of isselicrinus groenlandicus wienberg rasmussen (rosen krantz 1970; henderson et al. 1976; kollmann & peel 1983; petersen & vedelsby 2000). a fish fauna is also represented in the concretions (bendix-almgreen 1969) and scleractinian corals and coccoliths have been reported from the volcaniclastic sandstones (floris 1972; jürgensen & mikkelsen 1974). some corals are attached to tuff clasts, indicating that this was their original substrate. their random orientation suggests, however, that they are redeposited, but as they are only slightly worn, transport distances must have been short (floris 1972). a sparse nannoplankton and dinocyst flora is present at most localities (piasecki et al. 1992; nøhr-hansen et al. 2002), including the mudstones at the base of the volcanic breccias of the vaigat formation along the south coast of nuussuaq. the trace fossils planolites isp. and helmin thopsis horizontalis have been recognised in the gane#1 and gank#1 cores (dam 1996b, c). depositional environment. the macrofossils and the dinocysts indicate a marine depositional environment. based on the height of the overlying foresets in the hyaloclastite breccias, the water depth can be estimated to have been up to 700 m (a.k. pedersen et al. 1993). the sandstones are interpreted as the deposits of waning stage highto low-density turbidity currents. however, tuff layers deposited from suspension have also been recorded (a.k. pedersen et al. 1989). volcanism started from eruption centres in the north-western and western part of the region erupting directly into a deep marine environment covering large parts of nuussuaq and northern disko. the fossiliferous volcaniclastic sandstones and tuffs represent the distal bottomsets of major hyaloclastite beds of the vaigat formation (a.k. pedersen 1985), the result of sediment gravity flows that transported volcanic detritus and fossils into deeper water. boundaries. the lower boundary is placed at the base of the first volcaniclastic sandstones or tuffs. this boundco sand cl si pb 0 50 100 150 200 250 300 m q ui ka vs ak f or m at io n va ig at f or m at io n eq al ul ik f m fig. 120. reference section of the eqalulik formation overlying the paatuutkløften member (quikavsak formation) at ippigaarsukkløften on southern nuussuaq (for location, see fig. 40). the upper part of the eqalulik formation comprises fine-grained, tuffaceous mudstones deposited at water depths up to 700 m, as demonstrated by the height of the foresets in the overlying and laterally equivalent hyaloclastite breccias. modified from dam & nøhr-hansen (2001). for legend, see plate 1. 137 ary is, however, difficult to place in the field since outcrops are generally very poor in these mudstone-domi nated successions. along the north coast of nuussuaq and in the tunorsuaq valley, a conspicuous change in the colour of bottom sediment in small streams draining the local lithologies has been observed to occur at a distinctive level high in the sections. in the lower part of the sections, the stream-bed sediments are ochreous whereas this colouring is not present at higher levels; this change may reflect the sulphide content in the mudstones of the two formations and thus aid localisation of the formation boundary. if so, the mudstones of the kangilia formation were deposited in a poorly oxygenated, deepwater environment in contrast to the eqalulik formation where more oxygenated conditions could have resulted from intermittent disturbance of bottom waters by progradation of volcanic breccias. the upper boundary is generally placed at the base of the west greenland basalt group, at the base of the major hyaloclastic foreset beds of the vaigat formation or at sills situated along the base of the vaigat formation (hald & pedersen 1975; a.k. pedersen et al. 1993). on southern nuussuaq, however, between ataa and ippigaar sukkløften (fig. 40), the marine conditions prevailing during deposition of the eqalulik formation were succeeded by lacustrine environments, and consequently the eqalulik formation is overlain by non-marine shales of the atanikerluk formation (a.k. pedersen et al. 1996; g.k. pedersen et al. 1998). geological age. the eqalulik formation is visibly diachronous, reflecting the progressive eastward progradation of the hyaloclastite fans which are very well exposed along the slopes overlooking the south coast of nuussuaq 250 200 150 110 hyaloclastites 400 350 300 250 m m vf f m c vc sand pbsiclay gank#1 va ig at f m a nn aa na a m b eq al ul ik f or m at io n ?k an gi lia f or m at io n eq al ul ik f or m at io n vf f m c vc sandsi pbclay fig. 121. well reference section of the eqalulik formation in the gank#1 well. for location, see fig. 65; for legend, see plate 1. 138 (a.k. pedersen et al. 1993). however, the biostratigraphic resolution is in most cases not good enough to resolve this diachronism. the macrofauna at the type locality and on central nuussuaq indicates an early danian age (rosenkrantz 1970; floris 1972; henderson et al. 1976). the coccolith assemblage has been related to the np3 nannoplankton zone by jürgensen & mikkelsen (1974); however, new dinocyst and nannoplankton data indicate an np4 – possibly base np5 zone age (latest danian – early selandian; nøhr-hansen et al. 2002). 40ar/39ar age determinations of the volcanic rocks of the anaanaa member (vaigat formation) overlying the eqalulik for mation yield an age of 60.7 ± 1.0 (recalculated from storey et al. 1998). correlation. on the basis of dinocyst markers, the strata in the type section can be correlated with the volcanic hyaloclastites of the anaanaa and naujánguit members of the vaigat formation in the north-western part of the basin (nøhr-hansen et al. 2002). the tuffs of the abraham member (see below) have a unique composition with graphite and native iron that allows them to be correlated with the volcanic asuk member (a.k. pedersen et al. 1989). subdivision. the volcaniclastic sediments in agatdalen and qilakitsoq will continue to be assigned to a separate member, the abraham member, because of their unique geochemical composition which permits correlation of the sediments with the volcanic asuk member of the vaigat formation. abraham member revised member history. the abraham member was established by rosen krantz (in: koch 1959 pp. 75–78) as the uppermost member of the agatdal formation, but is now included in the eqalulik formation on the basis of the characteristic content of volcaniclastic material. name. after abraham løvstrøm from niaqornat, who for 30 years was a member of rosenkrantz’s expeditions to nuussuaq (fig. 7). distribution. the abraham member is known from two localities in the northernmost part of the agatdalen valley (fig. 113) and from the qilakitsoq area in the aaffarsuaq valley (fig. 82). type section. the type section of the abraham member is in turritellakløft at the northern end of agatdalen (fig. 113, 114, plate 3, section 1). the type section is located at 70°35.01´n, 53°08.02´w. reference section. a reference section occurs in ‘ravine 4’ east of qilakitsoq (figs 82, 83). thickness. the thickness of the abraham member is approximately 12 m in the type section, 10 m in qaersutjægerdal in agatdalen and 10–20 m at qilakitsoq. lithology. the abraham member consists of black and grey sandy mudstones intercalated with fossiliferous reworked volcaniclastic sandstones and basement pebble conglomerates (figs 114, 115, plate 3). volcanic material may constitute up to 25% of the section (a.k. fig. 122. fossiliferous, volcaniclastic sandstones of the eqalulik formation at danienrygge. pencil (encircled) and person for scale. 139 pedersen 1978). the beds are usually structureless, but cross-bedding has been recognised locally. volcaniclastic sandstones also occur as thin streaks in the mudstones and may include spherules of volcanic glass or tuff with a distinct chemical composition known only from the asuk member erupted from the ilugissoq graphite andesite volcano (a.k. pedersen 1985; a.k. pedersen & larsen 2006). in the lower part of the member, the mudstones are very dark, fossiliferous and rich in concretions. at one horizon, in situ corals are present in small bioherms, 10–20 cm high and less than 1 m wide (plate 3, section 4). this horizon can be followed for more than a kilometre on the eastern bank of qaersutjægerdal. the content of volcaniclastic sandstones increases upwards in the section. fossils. a sparse but diverse marine fauna and flora has been recorded from the abraham member, comprising scleractinian corals, echinoids, bivalves, gastropods, crustaceans, fish remains and palynomorphs (bendixalmgreen 1969; rosenkrantz 1970; floris 1972; koll mann & peel 1983; piasecki et al. 1992; petersen & vedelsby 2000). depositional environment. the macrofossil content in the lower part of the member indicates a marine depositional environment. however, an increase up-section in terrestrially derived palynomorphs indicates that the environment became increasingly brackish with time. sedimentary structures in the volcaniclastic sandstones indicate deposition from turbidity currents. the corals suggest a shallow water (50?–80 m) marine environment in a warm temperate climate (floris 1972), but where the corals were redeposited the water depth may well have exceeded 80 m. boundaries. a sharp lithological boundary occurs between the sandstones of the agatdal formation and the interbedded mudstones and volcaniclastic sandstones of the abraham member (figs 114, 115). at qilakitsoq and in the northern part of agatdalen, the abraham member is succeeded by hyaloclastites of the naujánguit member, whereas in the eastern part of agatdalen the abraham member is succeeded by hyaloclastites of the tunoqqu member (a.k. pedersen 1978; piasecki et al. 1992; a.k. pedersen, personal communication 1999). geological age. as for the eqalulik formation, i.e. latest danian to early selandian. correlation. the geochemical correlation of the abraham member with the hyaloclastites of the asuk member of the vaigat formation (a.k. pedersen 1978, 1985; a.k. pedersen & larsen 2006) ties these isolated outcrops to the mapped volcanic succession on the south coast of nuussuaq (a.k. pedersen et al. 1993). atanikerluk formation new formation history. the atanikerluk formation comprises the synvolcanic non-marine sediments in the nuussuaq basin; in the eastern part of the basin, it includes almost all the sediments overlying the atane formation (fig. 16). it is divided into five members (fig. 123), which are correlated to members in the volcanic vaigat and maligât formations (fig. 131). the intraand post-volcanic sediments of the nuussuaq basin are not included in the nuussuaq group and are therefore not discussed here. koch 1959 nuussuaqnuussuaq disko this paper point 976 member aussivik member umiussat member umiussat member umiussat member atanikerluk formation upper atanikerdluk formation pingu member naujât mb akunneq member assoq member assoq member naujât member naujât member quikavsak member not presentquikavsak formation fig. 123. lithostratigraphical subdivisions of the synvolcanic, non-marine deposits of the nuussuaq basin. 140 t 69°30’ 69°45’ 52°52°30’53° ku uk q aamasoq tuapaat ujarasussuk nuugaarsuk sullorsuaq vaigat aqajarua mudderbugten illukunnguaq sorte hak daug aar d-je nse n d al assoq niuluut illunnguaq siniffik marraat nuuk killuusat killu ssa ats ut fl ak ke rh uk innaarsuit skansen aamaruutissat skansentuapaat qaqqaat laksedalen killussaatsut kuuat akuliarutsip qaqqaa skorstensfjeldet blåbærdalen kuuk q aam asoq g ul e ry g fre de rik la ng e d al sullorsuaq kvandalen aqajaruata qaqqaa a ku nn eq inngigissoq pingu 5 km q e q e r t a r s u a q d i s k o se rm er su aq st or br æ en 500 maligât formation intrusion atanikerluk formation sandstone atanikerluk formation mudstone quaternary cover contour in metres atane formation skansen member ice sea/lake innanguit fig. 124. geological map showing the distribution of the atane and atanikerluk formations on eastern disko, simplified from a.k. pedersen et al. (2001). note that the outcrops of the atanikerluk formation are discontinuous. contour interval is 200 m. for location, see fig. 2. 141 in detail, however, the boundary between the atanikerluk formation and the volcanic formations can be complex, hyaloclastite breccias and invasive lavas interdigitating with the sedimentary succession. these sediments, interstratified locally with the lowermost volcanic layers, are intimately associated with the uppermost nuussuaq group strata and are thus referred to the atanikerluk formation where demonstrably related to an interdigitating volcanic–sediment facies front (figs 17, 131). two new formations, the quikavsak formation and the atanikerluk formation, replace the ‘upper atani kerd luk formation’ of nordenskiöld (1871), troelsen (1956) and koch (1959). the term ‘upper atanikerdluk for mation’ dates back to nordenskiöld (1871), who used this name for the beds (öfre atanekerdluklagren) containing the upper flora in the atanikerluk area. the plant fossils occur in concretions that are dark grey on fresh surfaces but weather to a dark red colour (nordenskiöld 1871 pp. 1051–52). koch (1959) reported that these concretions contain the upper atanikerdluk a flora of heer (1883a, b) and that they occur in his quikavsak member, i.e. the quikavsak formation of the present paper. the name atanikerluk formation is proposed because it does not imply the existence of a ‘lower atanikerd luk formation’. although a lower atanikerdluk flora was mentioned by heer (1868), this originated from the atane formation as recognised as early as 1871 by nordenskiöld. koch (1959) subdivided the upper atanikerdluk formation into five members on nuussuaq (fig. 123) but attempted no subdivision of the formation on disko. his quikavsak member is here established as the qui kavsak formation. two of koch’s members on nuus suaq are retained (the naujât and umiussat members), whereas the aussivik and point 976 members are abandoned and the sediments are referred to the new assoq member. the reason for this is that the aussivik and point 976 members are only found in a very small area of south-eastern nuussuaq (a.k. pedersen et al. 2007b). on disko, a subdivision into five members is proposed (the naujât, akunneq, pingu, umiussat, and assoq mem bers; fig. 123). the subdivision of the atanikerluk formation into five members reflects the dramatic palaeogeographic changes that accompanied the intense volcanic activity. name. the formation is named after the atanikerluk peninsula formed by a sill on the south coast of nuussuaq (fig. 40). the name was formerly spelled atanekerdluk or atanikerdluk. distribution. the atanikerluk formation is known on south-east nuussuaq (saqqaqdalen, atanikerluk, kingit toq, paatuut, and ataata kuua), on central nuussuaq (nassaat and locally in aaffarsuaq), on eastern disko (assoq, tuapaat qaqqaat, gule ryg, pingu, nuugaarsuk, frederik lange dal) and central disko (daugaard jensen dal and sorte hak; figs 2, 40, 124). in the atanikerluk area, the formation was mapped in great detail by koch & pedersen (1960). paleocene siliciclastic sediments are also present on svartenhuk halvø, where they are overlain by the volcanic vaigat and svartenhuk formations (j.g. larsen, personal communication 2008). these sediments remain unstudied in detail and the distribution of marine pre-volcanic, marine synvolcanic and nonmarine synvolcanic deposits is not known at present. in the future, paleocene non-marine, synvolcanic sediments on svartenhuk halvø may be established as one or more members within the atanikerluk formation. type section. the section in the south-facing slope between atanikerluk and tartunaq on south-east nuussuaq is retained as the type section (figs 125, 126) for the atanikerluk formation. this was described in detail by koch (1959), who defined type sections for several members of the formation in this area. the type section is located at 70° 03.63´n, 52°13.53´w. reference sections. reference sections of the atanikerluk formation are found at kingittoq, at pingu and in the tuapaat area (figs 7, 132, 140). the stratigraphic succession in north-east disko has been compiled from the (mainly sedimentary) sections on both sides of akunneq and the volcanic succession at aqajaruata qaqqaa, south of akunneq (a.k. pedersen et al. 2005) (figs 124, 131). thickness. the atanikerluk formation is up to 500 m thick in a composite section, but individual outcrops reach thicknesses of c. 400 m (pingu), c. 300 m (kingittoq, atanikerluk) and 200–250 m (saqqaqdalen) (figs 126, 128). at tuapaat qaqqaat, the thickness is difficult to measure due to landslides and interbedding of sediments and invasive lava flows (fig. 140). invasive lava flows are discussed below under ‘lithology’. lithology. the atanikerluk formation comprises mudstones, heterolithic sandstones, and fine-grained, loosely cemented sandstones. mudstones are dominant in the naujât and pingu members (fig. 127) and in the lower part of the assoq member, whereas very friable sandstones characterise the akunneq and umiussat members and the upper part of the assoq member. on a regional 142 scale, the formation comprises two coarsening-upward successions that may include thinner, coarsening-upward successions on a local scale (figs 16, 131). the mudstones are grey to dark grey, with abundant silt-sized particles and with kaolinite and quartz as the dominant minerals. in addition, gibbsite in subordinate amounts has been detected consistently in samples from the pingu member and commonly in the assoq member. locally, rows of small, yellowish brown siderite concretions occur in the mudstones (g.k. pedersen 1989; g.k. pedersen et al. 1998). reworked cretaceous palynomorphs are found in the lowest part of the atanikerluk formation at akunneq (hjortkjær 1991), suggesting that the formation also contains redeposited silt and clay. interbedded with the mudstones are thin layers of tuff. the particles are in the coarse sand fraction and have a brownish colour that weathers to pale buff or white. the strong alteration precludes a determination of the original chemical composition of the volcanic glass. the sandstones are weakly cemented, and in several outcrops the lithology may be better described as sand. colours range from pale grey to deep yellow, and fine coal debris is common. the sand or sandstones are generally fineto medium-grained; coarser grain sizes occur but are volumetrically insignificant. chert pebbles derived from ordovician sediments occur in the coarser facies, as also seen locally in the atane formation. the sediments of the atanikerluk formation could therefore include material reworked from cretaceous atane formation sand or sandstones. this is supported by the occurrences of few paleocene palynomorphs together with more numerous cretaceous spores and pollen in the lower part of the akunneq member at pingu (fig. 132, 285–355 m). the siliciclastic sediments are interbedded with volcanic rocks such as hyaloclastite breccias, invasive or subaqueous lava flows and tuffs. hyaloclastite breccias comprise particles ranging from boulder-sized pillow fragments to sand-sized grains of glass; these rocks formed qallorsuaq ‘keglen’ iviangernat assoq member invasive lava umiussat member atanikerluk formation naujât member atane formation kingittoq member subaerial lava flow fig. 125. the coastal slope above atanikerluk where koch (1959) described the strata of the atanikerluk formation (for location, see fig. 40). the peak at iviangernat is 1031 m a.s.l. the albian to cenomanian kingittoq member of the atane formation is up to 500 m thick and is overlain by the paleocene atanikerluk formation (see figs 16, 17). 143 when lava flowed into water. the breccias are often foreset-bedded and the height of the foresets provides an indication of water depth (jones & nelson 1970; a.k. pedersen et al. 1996; g.k. pedersen et al. 1998). invasive or subaqueous lava flows formed from large volumes of lava that reached the lake floor and continued into (invaded) the unconsolidated sediment without forming breccias (schmincke 1967; duffield et al. 1986; l.m. larsen & pedersen 1990; tucker & scott 2009). invasive or subaqueous lava flows can be traced laterally into sub20 30 clay silt sand vf f m c vc clay silt sand vf f m c vc 0 10 60 70 80 90 n au jâ t m em be r a ta ni ke rl uk f or m at io n a ta ni ke rl uk f or m at io n 100 110 120 130 140 150 160 170 180 m 0 10 20 10 20 30 0 m 0 10 20 30 0 40 50 0 10 20 30 0 40 50 40 50 t 2,4,5 t 5 t 3,1,2 ? u m iu ss at m em be r n au jâ t m em be r a ss oq m em be r * * quikavsak formation 1 3 2 fig. 126. composite type section of the atanikerluk formation from the atanikerluk area between naajaat and umiusat, south-eastern nuussuaq (for location, see fig. 40). 1: section through the naujât member from saqqaqdalen at the type locality defined by koch (1959) (fig. 129). 2: sections through the umiussat member from discontinuous outcrops in the umiusat area (figs 136, 137). 3: section through the upper part of the naujât member and the upper part of the assoq member, formerly the point 976 member of koch (1959) from discontinuous outcrops below keglen (fig. 125). for legend, see plate 1. 144 aerial lava flows and are thus different from sills, and are only slightly younger than the sediments they invade. many of the invasive lava flows have retained their chemical composition and may thus be correlated to the volcanic lithostratigraphy. layers of tuff are composed of sand-sized particles, and typically are less than 3 cm thick. the volcanic glass is strongly altered, and it is rarely possible to determine the original chemical composition of the glass. fossils. the atanikerluk formation contains a rich macro flora (koch 1959, 1963, 1964), which koch referred to the paleocene macclintockia zone characterised by meta sequoia occidentalis, cercidiphyllum arcticum, macclin tockia kanei, macclintockia lyalli and dicotylophyllum bellum (koch 1959, 1963). koch identified the upper atanikerdluk a flora and the upper atanikerdluk b flora of heer (1883a, b) in the basal part of the formation (lowermost naujât member). the upper atanikerdluk a flora is also present in the nuuk qiterleq member of the quikavsak formation. the difference between the two floras reflects changes in depositional environments, but the floras are essentially coeval (koch 1963). the diagnostic species of the macclintockia zone occur in the naujât member but have also been found in the abraham member of the eqalulik formation (koch 1963 p. 112). the palynomorphs (mainly spores and pollen) of the atanikerluk formation have been described by croxton (1978a, b), hjortkjær (1991), l.m. larsen et al. (1992), piasecki et al. (1992), lanstorp (1999) and d.j. mcintyre, personal communication 2009. marine dinocysts and rare bivalves have been reported locally from the assoq member. depositional environment. the mudstones are interpreted as lacustrine deposits on the basis of the high c/s ratios, the absence of pyrite, the presence of terrestrial fossils (macrofossil plants, spores and pollen) and the general absence of marine dinocysts (hjortkjær 1991; piasecki et al. 1992; g.k. pedersen et al. 1998). the sandstones are interpreted as dominantly fluvial and lacustrine in origin on the basis of the current-generated sedimentary structures, the lack of marine trace fossils, and their stratigraphic position relative to the lacustrine mudstones. periodic marine inundations are indicated by the presence of marine fossils in the mudstones of the assoq member (piasecki et al. 1992). boundaries. over large areas, the atanikerluk formation rests on an unconformity that defines the upper boundary of the atane formation and which reflects a hiatus of varying duration. in more restricted areas, the atani kerluk formation conformably overlies the quikavsak and eqalulik formations, or the volcanic vaigat formation (fig. 16). on south-eastern nuussuaq, the atanikerluk formation overlies either the atane or the quikavsak formation. strata of the atane formation dip towards the north-east, whereas the strata of the atanikerluk formation are sub-horizontal (fig. 129). where the quikavsak formation is present, its upper boundary is also horizontal, indicating that the angular unconformity between the atane and atanikerluk formations formed prior to the danian. at other outcrops, the lacustrine mudstones of the naujât member drape erosional relief at the top of the atane formation (figs 127, 130; pulvertaft 1989a; a.k. pedersen et al. 2007a). the lower boundary of the atanikerluk formation on nuussuaq is naujât member sill atane formation fig. 127. reference locality of the naujât member (atanikerluk formation) at kingittoq where lacustrine mudstones overlie the atane formation (see also g.k. pedersen et al. 1998 fig. 7). for location, see fig. 2. 145 10 0 20 30 40 50 60 70 80 90 110 100 120 130 150 160 170 180 190 200 210 220 230 m 140 clay silt sand pebbles vf f m c vc f m c clay silt sand pebbles vf f m c vc f m c a ku nn eq m em be r (u pp er p ar t) a ku nn eq m em be r (lo w er p ar t) pi ng u m em be r sk an se n m em be r a ta ni ke rl uk f or m at io n a ta ni ke rl uk f or m at io n a ta ni ke rl uk f or m at io n a ta ne f or m at io n 60 70 80 90 100 110 120 130 150 160 170 180 190 200 210 140 clay silt sand vf f m c vc pi ng u m em be r u m iu ss at m em be r 220 230 240 m a ku nn eq m em be r 50 30 25 3 6 30 10 5 15 4 15 6 6 8t t t t t t t t t t concretion 10 0 20 30 40 50 a ss oq m em be r 240 250 260 280 290 300 310 320 330 340 m 270 fig. 128. type section of the akunneq and pingu members (atanikerluk formation) measured at two closely spaced outcrops at pingu on eastern disko (for outcrop, see fig. 132; for location, see fig. 124). the boundary between the lower and upper parts of the akunneq member, at 110 m on the log (left), is at 335 m a.s.l. (fig 132). the locality is also the reference section for the atanikerluk formation, the umiussat member and the assoq member. for legend, see plate 1. 146 interpreted as a lacustrine flooding surface which can be mapped westwards into the volcanic terrain where it separates subaerial lavas of the naujánguit member from overlying hyaloclastite breccias at or just above the base of the ordlingassoq member (a.k. pedersen 1985; a.k. pedersen et al. 1993; g.k. pedersen et al. 1998) (fig. 131). on eastern disko, the lowest member of the atanikerluk formation is the dominantly fluvial akunneq member, and the base of this member constitutes the lower boundary of the formation; this is well exposed in the coastal section from pingu to nuu gaarsuk (figs 132, 133). the upper boundary of the atanikerluk formation is either a recent erosion surface or the boundary with volcanic or volcaniclastic deposits of the vaigat or maligât formations (figs 16, 131). geological age. the atanikerluk formation is dated by dinocysts, pollen or plant macrofossils, and is constrained by magnetostratigraphic and radiometric dating of the correlative volcanic units. a maximum age for the atanikerluk formation is obtained from the np4 – possibly early np5 marine dinocysts in the underlying eqalulik formation (nøhr-hansen et al. 2002). samples from the volcanic ordlingassoq and rinks dal members give radiometric ages of 60.7–61.1 ± 0.5–1.0 ma, recalculated from storey et al. (1998). this indicates an early to mid-paleocene (possibly selandian) age for the atanikerluk formation according to the timescale of gradstein et al. (2004) and ogg et al. (2008). correlation. the two regional, coarsening-upward successions recognised in the atanikerluk formation correlate with the volcanic ordlingassoq and rinks dal members of the vaigat and maligât formations respectively (fig. 131). correlation of individual members of the atanikerluk formation to strata of the vaigat for mation on disko presents some difficulties due to lack of exposures along the north coast of the island. the atanikerluk formation is younger than the eqalulik formation despite the similarity in macroplant fossils between the abraham and the naujât members. subdivision. the atanikerluk formation is subdivided into five members: the naujât, akunneq, pingu, umiussat and assoq members (fig. 123). naujât member revised member history. the naujât member was established by koch (1959). its distribution is expanded here, but otherwise the definition of the member is retained. name. the member is named from a small cove (modern spelling naajaat) on the south coast of nuussuaq, just west of saqqaqdalen (fig. 40; koch 1955). distribution. the distribution of the naujât member along the south-east coast of nuussuaq was mapped by koch (1959 plates 5–7). he recognised the naujât member in almost continuous outcrops and scree-covered slopes on the western side of saqqaqdalen and along naujât member atane formation s 70 m s q u fig. 129. type locality of the naujât member (atanikerluk formation) along the western slope of saqqaqdalen near naajaat (for location, see fig. 40). an angular unconformity is seen between the atane formation and the quikavsak formation (q). the latter is conformably overlain by lacustrine mudstones of the naujât member, which is cut by a sill (s). the outcrop of pale yellow sandstones in the distance belongs to the umiussat member (u). 147 the coast from naajaat to kingittoq (fig. 40). the member was traced in scree and landslides in the paatuut area, and, at ataata kuua, in a thin succession between the mudstones of the eqalulik formation and the hyaloclastite breccias of the ordlingassoq member below point 1010 m (figs 15, 16; a.k. pedersen et al. 1993, 2007b). in saqqaqdalen, the naujât member continues northwards for about 25 km along the upper western slopes of the valley (a.k. pedersen et al. 2007a). on disko, the naujât member is known from scattered outcrops along the north coast at qorlortorsuaq and towards nuugaarsuk (g.k. pedersen et al. 1998 fig. 12). the delineation of the member on disko is also discussed under the distribution of the pingu member. type section. the type section was described in the southwestern end of saqqaqdalen, above naajaat, by koch (1959; figs 124, 129). the type section is located at 70°04.10´n, 52°10.78´w. reference section. a reference section is proposed at kingittoq (fig. 127). a sedimentological log is shown in g.k. pedersen et al. (1998 fig. 10). thickness. the naujât member is thickest on nuussuaq, up to 230 m at kingittoq (koch 1959) and in saqqaqdalen (g.k. pedersen et al. 1998). on disko, the member is thin; up to 10 m are preserved at qorlortorsuaq on the north coast between qullissat and asuk (a.k. pedersen 1985). lithology. the naujât member comprises dark grey to black mudstones with thin, discontinuous layers of sandstones and thin tuff beds. the tuffs are strongly altered and the original chemical composition cannot be determined, but examination of thin sections suggests that the tuffs correlate with the vaigat formation. mineralogically, the mudstones are dominated by kaolinite and quartz with minor feldspar, illite and smectite. gibbsite has only been identified in 15% of the samples, mostly from the upper part of the member. pyrite has not been detected. atane formation, kingittoq member naujât member fig. 130. outcrop of the naujât member overlying the atane formation at eqip inaarsuata qaqqaa, at the northern end of the saqqaqdalen valley (located as eiq in fig. 2). the lacustrine mudstones drape erosional relief on the top of the atane formation. the naujât member is c. 200 m thick. 148 lower rinks dal mb upper rinks dal mb ordlingassoq mb naujánguit mb 60.8±0.5 60.7±1.0 60.9±0.4 60.7±0.4 anaanaa mb va ig at f m m al ig ât f m akuarut unit skarvefjeld unit 61.1±0.5 assoq mbassoq mb assoq mb akunneq mb a ta ni ke rl uk fo rm at io n a ta ni ke rl uk fo rm at io n condensed section subaerial surface umiussat mb naujât mb umiussat mb pingu mb skarvefjeld 5 assoq 6 tuapaat 7 atanikerluk 4 giesecke m.paatuut 1 fr. langes dal peak 1123 m 8 ingigissoq 10 pingu 11 marraat kingittoq lower rinks dal mb upper rinks dal mb ordlingassoq mb naujánguit mb anaanaa mb va ig at f m m al ig ât f m akuarut unit skarvefjeld unit disko 70°n 53°w hareøen 50 km5 6 7 9 8 11 10 4 3 1 2 condensed section naujât mb qullissat inussuk fjeld2 3 9 eqalulik fm subaerial lava flow subaqueous lava flow columnar jointing entablature zone hyaloclastite breccia sandstone lacustrine mudstone marine mudstone conglomerate coal marine dinoflagellate nuussuaq fig. 131. simplified stratigraphic sections showing the relationships between the volcanic rocks (colour coded) of the vaigat and maligât formations (west greenland basalt group) and the atanikerluk formation. the upper boundary of the assoq member is indicated (dashed line). the invasive, subaqueous lavas are all assigned to the maligât formation. the radiometric ages are recalculated from storey et al. (1998). 149 the toc [total organic carbon] content is high (up to 11%) and c/s ratios are high (g.k. pedersen et al. 1998). perregaard & schiener (1979) noted that the organic matter is immature, consisting of exinite (c. 50%), vitrinite (c. 35%) and inertinite (c. 15%), and that it is chemically dominated by saturated and aromatic hydrocarbons. a study of palynofacies indicates a predominance of brown and black lignite (hjortkjær 1991). fossils. the naujât member has yielded well-preserved macroplant fossils, the upper atanikerdluk flora a and b of heer (1883a, b; koch 1959, 1963). leaves from deciduous trees are an important constituent in the flora which includes: cladophlebis groenlandica, metasequoia occidentalis, cercidiphyllum arcticum, dicotylophyllum bellum, dicotylophyllum steenstrupianum, macclintokia kanei, macclintockia lyalli and credneria spectabilis (koch 1963). the assemblage of spores and pollen in the naujât member was studied by hjortkjær (1991). the flora is dominated by palynomorphs of terrestrial origin, whereas remains of lacustrine plants and algae are rare. pollen of taxodium spp. are abundant. the stratigraphically important mid-paleocene species momipites actinus and carya pollenites wodehouseia are present in this member. depositional environment. the naujât member is interpreted as a succession of lacustrine mudstones (koch 1959; schiener & leythaeuser, 1978; g.k. pedersen et al. 1998). the lake formed through damming by the volcanic rocks of the ordlingassoq member, and it was filled simultaneously by hyaloclastite breccias from the west and siliciclastic mud from the east and south-east (g.k. pedersen et al. 1998). the abundance of kaolinite among the clay minerals suggests that these have the same provenance as the mudstones of the cretaceous atane formation, probably areas of deeply weathered crystalline rocks east of the basin boundary fault. boundaries. the lower boundary of the naujât member corresponds to the lower boundary of the atanikerluk formation in the area where the naujât member occurs (figs 126, 129, 130). mapping of the lower boundary is difficult in the paatuut and ataata kuua areas where the member overlies marine mudstones of the eqalulik formation. in this area, the naujât member is also interbedded with the toesets of hyaloclastite breccias of the ordlingassoq member (fig. 17, locality 5). the upper boundary of the naujât member is everywhere towards either the umiussat member or the ordlingassoq member of the vaigat formation (fig. 136). geological age. within the resolution of the dating methods, the age of the naujât member lies within the age range of the atanikerluk formation (i.e. early to midpaleocene, see p. 146). correlation. the naujât member is coeval with the volcanic breccias of the ordlingassoq member; both members overlie the same lacustrine flooding surface (a.k. pedersen et al. 1993; g.k. pedersen et al. 1998). the lacustrine naujât member on nuussuaq correlates with the fluvial akunneq member on north-eastern disko. the palynomorph assemblages of the naujât, akunneq and pingu members are similar, and the upper part of the naujât member correlates with the pingu member (figs 123, 131). akunneq member new member history. the sediments now referred to the akunneq member were formerly referred to the upper atanikerdluk formation (koch 1964; croxton 1978a section c2). name. the member is named after the akunneq valley between the pingu and inngigissoq mountains on the north-east coast of disko. distribution. the akunneq member is known from the north-east coast of disko (pingu to nuugaarsuk) and southwards to gule ryg (fig. 124). type section. the section at pingu on the eastern side of akunneq is chosen as the type section (figs 128, 132). the type section is located at 69°47.53´n, 52°05.43´w. reference section. the section on the western side of akunneq is chosen as the reference section. thickness. the akunneq member is c.165 m thick at pingu but only c.145 m at nuugaarsuk. thus the member decreases in thickness towards the west (fig. 133). lithology. the akunneq member is dominated by very friable sandstone. it comprises a coarse-grained lower part and a finer-grained upper part, separated by a mudstone horizon at 335 m a.s.l. in the pingu–akunneq sec150 tion (figs 132, 133). the lower part consists of large-scale cross-beddded, coarse-grained sandstones that differ little from the fluvial sandstones of the atane formation (figs 128, 134). the upper part is dominated by finegrained white sands interbedded with thinner horizons of mudstone, carbonaceous mudstone or thin coal seams. bedding planes in the sandstone are often draped by coaly plant debris. photogrammetric work shows that the c. 5 m thick mudstones are laterally continuous over a distance of 10 km, and that the proportion of finegrained facies increases in a westerly direction from pingu to nuugaarsuk. locally, the upper sandstones include 2–3 m thick beds with low-angle, composite cross-bedding. bedform migration directions suggest westerly palaeocurrents. fossils. thin mudstone horizons in the akunneq member contain paleocene spores and pollen as well as reworked species of cretaceous (cenomanian) age (croxton 1978a, b; hjortkjær 1991, d.j. mcintyre, personal communication 2009). the upper part of the akunneq member is dominated by paleocene species identical to those occurring in the naujât and pingu members (hjortkjær 1991; b.f. hjortkjær, personal communication 1999). the lower part of the akunneq member is dominated by reworked cenomanian species, and unquestionable paleocene species are scarce (b.f. hjortkjær, personal communication 1999; d.j. mcintyre, personal communication 2009). depositional environment. the akunneq member is interpreted as a succession of sandy fluvial deposits interbedded with thin lacustrine mudstones. the sandstones in the lower part of the akunneq member are interpreted as deposited in braided channels. the change in grainsize and the ubiquitous coaly plant debris indicate lower energy, and the occasional occurrence of sandy point bar deposits suggests deposition in meandering fluvial channels. the thin but laterally widespread mudstones are interpreted as reflecting short-lived phases of lake formation, and the coal beds represent peat formation. towards the west, in a downstream direction, the akunneq upper rinks dal member assoq member pingu 845 m akunneq member atane formation skansen member umiussat member pingu member akunneq member sill sill sill sill inngigissoq 845 m u. rinks dal member p d d 335 m a.s.l. 285 m a.s.l. fig. 132. outcrop of atanikerluk formation in the area between pingu and inngigissoq on eastern disko cut by several dykes (d) and sills (for location, see fig. 124). the boundary between the atane formation (skansen member) and the atanikerluk formation is set at 285 m a.s.l. the boundary at 335 m a.s.l. separates the lower and upper parts of the akunneq member (for section, see fig. 128). p, pingu member. 151 600 500 400 300 220 600 m 500 420 400 300 260 mud sand 700 pingu a ss oq m em be r u m iu ss at m em be r pi ng u m em be r a ku nn eq m em be r sk an se n m em be r a ta ne f or m at io n a ta ni ke rl uk f or m at io n nuugaarsuk m ? ? mud sand fig. 133. correlation of the members within the atanikerluk formation on north-east disko. note the general thinning of units westwards from pingu to nuugaarsuk. altitudes above sea level are indicated. for location, see fig. 124; for legend, see plate 1. 152 member becomes thinner with an increasing amount of mudstone, which suggests that it was deposited on a floodplain adjacent to the ‘naujât lake’. boundaries. the lower boundary of the akunneq member corresponds to the lower boundary of the atanikerluk formation in the area of distribution of the akunneq member (fig. 132). the geological map of the pingu area shows that both the atane and the atanikerluk formations are subhorizontal in north-east disko (a.k. pedersen et al. 2001). the boundary between the fluvial skansen member and the fluvial akunneq member may be difficult to identify in outcrops. in the pingu area, the lower boundary of the akunneq member is placed at the base of a c. 3 m thick unit of bluish grey, laterally continuous mudstones, which forms a terrace with a steep front. the upper boundary of the akunneq member is placed at the base of the mudstones of the pingu member (figs 132, 133, 135). geological age. reworked cretaceous spores and pollen dominate the lower part of the akunneq member, but the few paleocene forms show that the age of the akunneq member lies within the age range of the atanikerluk formation (i.e. early to mid-paleocene, see above). correlation. the akunneq member is interpreted to correlate with the lower or middle part of the naujât member. abundant landslides on the north coast of disko obscure the relationships between the akunneq member and the ordlingassoq member of the vaigat formation (a.k. pedersen et al. 2005). pingu member new member history. sediments referred to the pingu member were earlier described by croxton (1978a section c2). their sedimentary facies and depositional environment were discussed by g.k. pedersen (1987, 1989). the deposits were tentatively referred to the naujât member by g.k. pedersen (1987). the difference in mineralogy, the lack of positive evidence that the pingu and naujât members are continuous, and the prominence of this mudstone unit on north-east disko are the reasons for erecting the pingu member. name. the member is named from the mountain of pingu on north-east disko (figs 124, 132). distribution. the pingu member is only known from outcrops on disko. it is continuously exposed in the coastal section from pingu to nuugaarsuk and can be traced south-west of pingu to gule ryg (fig. 124). type section. the section on the north side of pingu (east of akunneq) is chosen as the type section of the pingu member (figs 128, 135). the type section is located at 69°47.13´n, 52°02.12´w. reference section. the section between akunneq and nuu gaarsuk is chosen as the reference section of the pingu member (figs 124, 133). fig. 134. the illustrated, coarse-grained, fluvial sandstone unit, c. 30 m thick, showing large-scale cross-bedding is characteristic of the lower part of the akunneq member on the western side of akunneq. the sandstone is abruptly overlain by grey muddy sandstones, also referred to the akunneq member. for location, see fig. 124. 153 thickness. the pingu member is 85 m thick at pingu, probably thinning towards nuugaarsuk, although the thickness is difficult to measure at the latter locality due to landslides (fig. 133). the member is more than 30 m thick at gule ryg. lithology. the pingu member consists of dark grey mudstones interbedded with thin layers of tuff and finegrained sandstone beds that increase in frequency upwards. the member thus constitutes an overall coarseningupward succession (fig. 128; g.k. pedersen 1989). mineralogically, the mudstones consist of kaolinite and quartz with a little illite and feldspar; neither calcite nor pyrite has been detected. gibbsite is found in 96% of the samples where it makes up 5–10% of the sediment. siderite occurs in small concretions that weather bright yellow. the mudstones are rich in organic matter (up to 8%), most of which is terrestrial (g.k. pedersen 1989). a palynofacies study demonstrated the predominance of brown and black wood (hjortkjær 1991). the sandstones are fine-grained, well-sorted and form thin beds showing parallel lamination or current ripple cross-lamination. fossils. the assemblage of spores and pollen in the pingu member was studied by hjortkjær (1991), who found that it corresponds to that of the naujât member. the flora is dominated by palynomorphs of terrestrial origin, whereas remains of lacustrine plants and algae are rare. pollen of taxodium spp. are abundant. the stratigraphically important mid-paleocene species momipites actinus and caryapollenites wodehouseia occur together with a few specimens of the midto late paleocene insula polle nites rugulatus (hjortkjær 1991). neither croxton (1978a, b) nor hjortkjær (1991) observed marine dinocysts in samples from the pingu member. depositional environment. the pingu member represents a predominantly low energy depositional environment characterised by settling of mud from suspension and occasional deposition of sand from low energy sediment gravity flows. the lack of marine fossils coupled with the absence of both pyrite and its weathering product jarosite indicates a lacustrine depositional environment. the increasing number of sand layers up-section are interpreted to record gradual progradation of the shoreline and consequent filling of the lake (g.k. pedersen 1989). the supply of gibbsite to the lacustrine mud suggests input from a new provenance area. the expanding areas of subaerial basalt flows may have weathered to lateritic soils that could have supplied gibbsite during deposition in the ‘pingu lake’. this new provenance area is also reflected in the mineralogy of the top of the lacustrine naujât member. boundaries. the pingu member has a sharp lower and a gradational upper boundary (figs 128, 132, 135). the lower boundary separates the lacustrine mudstones of the pingu member from the underlying sandy fluvial akunneq member; this abrupt boundary is overlain either directly by mudstones or by a thin succession of mudstones interbedded with sandstone layers and it is interpreted as an erosional surface formed during lacustrine drowning (figs 128, 135). the upper boundary is defined by the base of the umiussat member (figs 128, 132, 133, 134, 135). pingu member umiussat member 25 m akunneq member fig. 135. type locality of the pingu member on the north side of pingu (for location, see fig. 124). note the sharp lower boundary with the akunneq member and the transitional upper boundary with the umiussat member. 154 geological age. the age of the pingu member is the same that of the atanikerluk formation (i.e. early to midpaleocene, see p. 146). correlation. the deposits assigned here to the pingu member have been correlated with the naujât member (henderson et al. 1981 fig. 4). the assemblages of spores and pollen in the two members are similar (hjortkjær 1991), but the mineralogy differs, especially with respect to the distribution of gibbsite. we interpret the pingu member on disko as correlating with the upper part of the naujât member on nuussuaq. a correlation between the pingu member and the volcanic ordlingassoq member cannot be proven because the north coast of disko is ravaged by landslides and rock glaciers between nuugaarsuk and qullissat (a.k. pedersen et al. 2005). between pingu and nuugaarsuk, the oldest magmatic rocks have a composition corresponding to the maligât formation (fig. 131). umiussat member revised member history. the umiussat member was established by koch (1959) on south-east nuussuaq where it overlies the naujât member. in the present paper the definition of the umiussat member is retained, but its geographical distribution is expanded. name. the member is named from the umiusat ridge on south-eastern nuussuaq (fig. 40). distribution. the umiussat member is known on southeast nuussuaq in the area between naajaat and kingittoq, and in saqqaqdalen. it covers a smaller area than the naujât member. the umiussat member is also present on eastern disko between nuugaarsuk and pingu, and at gule ryg (figs 40, 124). type section. the type section of the umiussat member of koch (1959) at umiusat is retained (figs 126, 136). the type section is located at 70°09.00´n, 52°26.57´w. reference section. the section on the north side of pingu, east of akunneq, is suggested as a reference section for the umiussat member (figs 128, 132). thickness. the umiussat member is 70−100 m thick. it is rarely well exposed, hence lateral variations in thickness are not documented. the member is c. 80 m thick at pingu (fig. 128) and c. 100 m thick in the atanikerluk area (koch 1959). lithology. the umiussat member is dominated by fineto medium-grained sands or sandstones which range in colour from white or pale grey on disko to yellow on nuussuaq. comminuted plant debris is abundant and outlines bedding planes and sedimentary structures such as cross-lamination and cross-bedding. the sands or sandstones are interbedded with 1–5 m thick mudstones, naujât member maligât formation lava flows umiussat member 35 m fig. 136. type locality of the umiussat member (atanikerluk formation) at umiusat, south-eastern nuussuaq (for location, see fig. 40). at the top of the naujât member, an increase in the number of thin sandstone beds results in a transitional boundary between the two members. 155 some of which include thin coal beds (figs 126, 128, 137). a 120 cm thick coal bed forms the top of the umiussat member on north-east disko (fig. 131). fossils. no fossils have been reported from the umiussat member. depositional environment. the sediments of the umiussat member are interpreted to have been deposited in predominantly braided fluvial channels and on floodplains. boundaries. the lower boundary with the naujât member is transitional and is rarely well exposed (fig. 136; koch 1959). the lower boundary with the pingu member is placed at an erosional surface separating the heterolithic sandy mudstones at the top of the pingu member from the overlying fluvial sandstones of the umiussat member (figs 128, 135). the upper boundary of the umiussat member is placed at the drowning surface which forms the lower boundary of the assoq member (figs 133, 137). geological age. the age of the member lies within the age range of the atanikerluk formation (i.e. early to midpaleocene, see p. 146). correlation. the umiussat member is interpreted as coeval with the uppermost subaerial lava flows of the ordlingassoq member, and deposition of the umiussat member probably continued during the break in volcanic activity between the vaigat and the maligât formations (fig. 131). assoq member new member history. the sediments referred to the assoq member include those overlying the umiussat member on disko as well as those referred to the aussivik member and the point 976 member of koch (1959) (fig. 123). the aussivik and point 976 members are abandoned because they are poorly exposed and only cover small areas on south-eastern nuussuaq. name. the member is named from the coastal mountain slope at assoq on southern disko (fig. 124). distribution. the geological map sheets 1:100 000 uiffaq and 1:100 000 pingu show that the assoq member covers most of disko, east of the disko gneiss ridge (a.k. pedersen et al. 2000, 2001). the area in which the assoq member is distributed is strongly affected by landslides and reliable outcrops are discontinuous (fig. 124). on nuussuaq, the assoq member is present at kingittoq, between umiusat and saqqaqdalen, in the atanikerluk area and probably also on central nuussuaq, east of the ikorfat fault (fig. 40; a.k. pedersen et al. 1993, 2002, 2007b). type section. the type section is at assoq within a huge landslipped block (fig. 138). this locality has the best exposures of the fissile mudstones, but neither the lower nor the upper boundary of the assoq member is confimaligât formation umiussat member assoq member15 m fig. 137. lower, sharp boundary of the assoq member with the umiussat member at umiusat. see section 2 in the composite sedimentological log in fig. 126. 156 dently identified (fig. 139). the type section is located at 69°19.23´n, 53°09.10´w. reference sections. reference sections are exposed on southern disko east of assoq (fig. 139), on northern disko at pingu (figs 124, 128), and on southern nuussuaq at qallorsuaq (keglen). the sediments here were formerly referred to the aussivik and point 976 members of koch (1959) (figs 40, 125). thickness. the assoq member is up to c. 200 m thick, but the thickness is difficult to measure in areas where the sediments are interbedded with volcanic or intrusive rocks and are subject to landslides. lithology. the assoq member constitutes a major, coarsening-upward succession with a lower part dominated by brownish black fissile mudstones, and an upper part that is dominated by sands but also includes a single, 3 m thick, coal seam (figs 126, 128, 138, 140). the mudstones have toc contents of up to 7%. the dominant clay mineral is kaolinite, but gibbsite is also present in most samples. thin beds are cemented by siderite. nume rous thin tuff layers are interbedded with the mudstones. they are typically a few millimetres to a few centimetres thick, deposited by settling through the water column. examination of thin sections shows that the tuff is strongly altered diagenetically. the mudstones are interbedded with hyaloclastite breccias and subaqueous lava flows which formed where lava flows of the lower rinks dal member entered the ‘assoq lake’ (l.m. larsen et al. 2006). based on their chemical composition, the invasive lava flows are correlated with various units within the volcanic lower rinks dal member (figs 131, 139, 140). assoq member invasive lava invasive lava invasive lava assoq member subaqueous lava subaqueous lava upper rinks dal memberlower rinks dal member subaqueous lava 55 m a.s.l. fig. 138. type locality of the assoq member at assoq on the south coast of disko (for location, see fig. 124). the brownish black mudstones of the assoq member are interbedded with invasive lava flows and overlain by subaqueous lava flows of the lower rinks dal member of the maligât formation (see fig. 131). sediments and interbedded volcanic rocks are part of a huge landslide, two blocks of which are separated by the dash–dot line. the assoq member is also found at higher altitudes away from the coast (a.k. pedersen et al. 2003). for log, see fig. 139. 157 the upper part of the assoq member is dominated by sands, often with comminuted plant debris. in the scree-covered slopes at tuapaat qaqqaat and qallorsuaq (keglen), the sands are yellow and fine-grained, but rarely well exposed. at pingu and gule ryg, the sedimentary structures suggest deposition from traction currents in a low-energy environment (fig. 128). a thick coal bed has been observed at the top of the assoq member (fig. 141) which probably correlates to coal beds interbedded with volcanic rocks at several localities west of gule ryg, eastern disko (l.m. larsen & pedersen 1990; a.k. pedersen et al. 2001) (fig. 131). these sediments are considered part of the atanikerluk formation as they can be traced from continuous sedimentary successions into the basal part of the subaerial lava flow succession. the chemistry of these volcanic rocks corresponds to the lower part of the upper rinks dal member. fossils. plant macrofossils are unevenly distributed in the assoq member, but they are generally scarce. from the atanikerluk area, koch (1959) reported a leaf of cercidiphyllum arcticum from the lower part (his aussivik member) and pieces of fossil wood of taxodioxylon type from his point 976 member. bedding planes covered by plant remains were observed west of kingittoq by the present authors. the freshwater bivalve unio sp. occurs in the upper, sandy part of the assoq member (koch 1959). fish scales and bones are preserved in a concretion from akuliarusinnguaq (c. 4 km north of gieseckes monument; fig. 2). the spores and pollen in the assoq member were described by hjortkjær (1991) and differ little from those in the older members of the atanikerluk formation. piasecki et al. (1992) reported finds of marine dinocysts in a few samples from the assoq member in southern and eastern disko. in the inner part of kvandalen (fig. 124), the mudstones are found to contain rare bivalves (protobranchs?) and rare dinocysts (piasecki et al. 1992). a few marine dinocysts were found at tuapaat qaqqaat at the top of the assoq member just below the basalt conglomerate (fig. 140). 0 5 10 25 30 35 invasive lava lower rinks dal mb invasive lava lower rinks dal mb invasive lava lower rinks dal mb sill 40 105 110 125 130 m t t clay si sand t t t t t t t t t t t t s t s s s s ?a ta ne f or m at io n a ta ni ke rl uk f or m at io n a ss oq m em be r 8 10 30 20 22 10 3 7 10 10 10 10 10 20 30 12 10 5 fig. 139. type section of the assoq member at assoq on the south coast of disko (for location, see fig. 124). the lacustrine mudstones contain numerous millimetre-thick tuff beds. for legend, see plate 1. depositional environment. at skarvefjeld (fig. 2), the lower, shaly part of the assoq member is interbedded with hyaloclastite breccias that are traced laterally into the lower rinks dal member (heinesen 1987; l.m. larsen & pedersen 1990). water depths of 80−100 m are calculated from the height of the foresets of the hyaloclastite breccias. similar interbedding of assoq member mudstones and hyaloclastite breccias is also known from sorte hak on central disko, slightly east of the disko gneiss ridge (fig. 124). the assoq member is interpreted as lacustrine in origin, deposited within the ‘assoq lake’. at present it is not possible to demonstrate whether this was one huge lake or several smaller lakes. scarce marine dinocysts indicate that the assoq lake was subject to marine inundations. the dinocysts were found in samples from eastern and south-eastern disko, but were not recorded from nuussuaq. this distribution, coupled with a south-eastward tilting of an originally nearly horizontal magmatic boundary (l.m. larsen & pedersen 1992) suggests that the marine inundations were from the south. the upward transition from mudstones to sandstones reflects shallowing and gradual fill of the lake. during a break in clastic sediment input, peat accumulated in a large swamp in eastern disko, and resulted in a thick coal bed (fig. 141). boundaries. the lower boundary of the assoq member is placed at the base of the mudstone succession, and is interpreted as a lacustrine drowning surface (figs 128, 137). on north-east disko, the lower boundary overlies a coal bed. the upper boundary is placed where the sediments are overlain by subaerial lava flows of the rinks dal member. note that locally thin tongues of sediment within the lower rinks dal member are assigned to the assoq member; such sediment wedges are bounded abruptly beneath and above by volcanic strata. at tuapaat qaqqaat, the uppermost bed in the assoq member is a conglomerate with clasts of basaltic rocks derived from lava flows of the akuarut unit of the rinks dal member (fig. 140; l.m. larsen & pedersen 1990, 2009). 0 m clay silt sand pebbles vf f m c vc f m c 280 250 200 150 100 50 380 400 450 500 510 covered by scree a ss oq m em be r sk an se n m em be r a ta ni ke rl uk f or m at io n a ta ne f or m at io n in va si ve la va fl ow s co rr el at e to t he lo w er r in ks d al m em be r in va si ve la va fl ow s co rr el at e to t he a ku ar ut u ni t of t he lo w er r in ks d al m em be r fig. 140. composite reference section of the assoq member at tuapaat qaqqaat, south-eastern disko (for location, see fig. 124). note the occurrence of marine dinocysts at the top of the lacustrine section, indicating a brief marine inundation. colour coding of volcanic rocks follows fig. 131. for legend, see plate 1. 158 159 geological age. the age of the assoq member lies within the age range of the atanikerluk formation (i.e. early to mid-paleocene, see p. 146). correlation. on southern disko, the lower, shaly, part of the assoq member is interbedded with subaqueous lava flows and hyaloclastite breccias, which continue westwards into the skarvefjeld unit (formerly pahoehoe unit) of the lower rinks dal member (fig. 131; heinesen 1987; l.m. larsen & pedersen 1990, 2009; a.k. pedersen et al. 2003; l.m. larsen et al. 2006). on north-eastern disko (akunneq and pingu), the earliest volcanic rocks are thick invasive lava flows with columnar jointing, which invaded the uppermost part of the umiussat member and the assoq member. the composition of these volcanic rocks tie them to the uppermost lower rinks dal member and the akuarut unit (formerly feti unit) (fig. 131; a.k. pedersen et al. 2005; l.m. larsen et al. 2006; l.m. larsen & pedersen 2009). on nuussuaq, lava flows of the akuarut unit invaded mudstones of the assoq member and increased markedly in thickness due to ponding of the lavas in the ‘assoq lake’. throughout eastern disko and southern nuussuaq the assoq member is thus interbedded with invasive flows of the rinks dal member, and the various volcanic facies (breccias, invasive lavas or ponded subaerial lava flows) are thought to reflect different extrusion rates. in most of its area of distribution, the assoq member is overlain by subaerial lava flows of the upper rinks dal member (fig. 131). it is concluded that the assoq member generally correlates with the rinks dal member. on easternmost central nuussuaq, however, an outcrop of mudstones interbedded with invasive lava flows of the younger niaqussat member is also referred to the assoq member (a.k. pedersen et al. 2002). acknowledgements the work reported on in this publication was carried out with considerable financial support from the danish energy research programme (efp), the danish natural science research council (snf), the bureau of minerals and petroleum, government of greenland, the carlsberg foundation and the arktisk station, university of copen hagen. their support is gratefully acknowledged. the authors are greatly indebted to the three referees: dr. t. christopher r. pulvertaft, now retired from geus, has taken a keen interest in the manuscript from its early stages and has encouraged the authors throughout. he has corrected the english language and has used his great knowledge of the area to ensure that details in descriptions and illustrations are correct. dr. j. christopher harrison of the geological survey of canada and dr. robert knox of the british geological survey have read the manuscript very carefully and offered a large number of helpful suggestions for the improvement of text and figures. with few exceptions, the illustrations have been prepared by jette halskov, geus, without whom the manuscript would not have been completed. upper rinks dal member coal assoq member fig. 141. upper part of the assoq member including a 5.1 m thick coal bed. south side of blåbærdalen, eastern disko (for location, see fig. 124). correlative coal beds are seen at several widely spaced outcrops indicating the development of large swamps. the sediments are overlain by subaerial lava flows of the volcanic rinks dal member (see fig. 131). ambirk, d. 2000: a sedimentary, sequence stratigraphic and carbon isotopic investigation of sections of the atane formation, nuussuaq, central west greenland, 137 pp. unpublished m.sc. thesis, university of copenhagen, denmark. balkwill, h.r., mcmillan, n.j., maclean, b., williams, g.l. & srivastava, s.p. 1990: geology of the labrador shelf, baffin bay, and davis strait. in: keen, m.j. & williams, g.l. 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and greenland bulletin 11, 185–204. yen, j.t.c. 1958: a cretaceous non-marine molluscan fauna of west greenland. bulletin grønlands geologiske undersøgelse 21, 13 pp. 168 a aaffarsuaq central nuussuaq figs 2, 82 aamaruutissat (also skansen) southern disko fig. 124 aasiaat southern disko bugt figs 2, 10 agatdalen central nuussuaq fig. 113 agatkløft central nuussuaq fig. 113 akuliarusinnguaq (ak) southern nuussuaq fig. 2 akunneq northern disko fig. 124 alianaatsunnguaq southern nuussuaq fig. 2 alianaitsúnguaq (now alianaatsunnguaq) southern nuussuaq figs 2, 4 anariatorfik western nuussuaq fig. 65 angiarsuit northern nuussuaq fig. 22 angiissat (part of grønne ejland) southern disko bugt fig. 2 angnertuneq (now annertuneq) annertuneq northern nuussuaq fig. 74 aqajaruata qaqqaa eastern disko fig. 124 auvfarssuaq (now aaffarsuaq) assoq southern disko figs 2, 124 asuk northern disko fig. 2 ata (now ataa) southern nuussuaq fig. 4 atâ (now ataa) ataa southern nuussuaq fig. 40 ataata kuua southern nuussuaq figs 6, 40 atane (now ataa) southern nuussuaq fig. 3 atanekerdluk (now atanikerluk) southern nuussuaq fig. 3 atanikerdluk (now atanikerluk) southern nuussuaq fig. 4 atanikerluk southern nuussuaq fig. 40 atâta kûa (now ataata kuua) b baffin bay fig. 1 blåbærdalen southern disko fig. 124 d danienrygge (dr) northern nuussuaq fig. 74 daugaard jensen dal central disko fig. 124 davis strait fig. 1 disko figs 2, 10, 124 disko bugt figs 2, 10 disko gneiss ridge (dgr) western disko fig. 10 e ekkorfat (now ikorfat) northern nuussuaq fig. 3 ekorgfat (now ikorfat) northern nuussuaq fig. 4 eqalulik western nuussuaq fig. 65 eqip inaarsuata qaqqaa (eiq) eastern nuussuaq fig. 2 f firefjeld svartenhuk halvø fig. 73 fp93-3-1 northern disko fig. 2 fp94-11-02 northern nuussuaq fig. 74 fp94-11-04 northern nuussuaq fig. 74 fp94-11-05 northern nuussuaq fig. 74 fp94-9-01 western nuussuaq fig. 65 frederik lange dal eastern disko fig. 124 g gane#1 western nuussuaq fig. 65 gank#1 western nuussuaq fig. 65 gant#1 northern nuussuaq fig. 74 ganw#1 western nuussuaq fig. 65 ggu 247701 southern nuussuaq fig. 40 ggu 247801 southern nuussuaq fig. 40 ggu 247901 southern nuussuaq fig. 40 ggu 400701 central nuussuaq fig. 113 ggu 400702 central nuussuaq fig. 113 ggu 400703 central nuussuaq fig. 113 ggu 400704 central nuussuaq fig. 113 ggu 400705 northern nuussuaq fig. 74 ggu 400706 northern nuussuaq fig. 74 ggu 400707 northern nuussuaq fig. 74 ggu 400708 svartenhuk halvø fig. 73 ggu 400709 svartenhuk halvø fig. 73 ggu 400710 svartenhuk halvø fig. 73 ggu 400711 svartenhuk halvø fig. 73 ggu 400712 svartenhuk halvø fig. 73 giesecke monument southern nuussuaq figs 6, 40 169 appendix: place names and localities the place names and localities mentioned in the text are shown on one or more maps. figure 2 contains many names and indicates the position of more detailed maps where the remaining names are shown. names in italics are names with old spelling. they are located under their modern spelling or indicated on the historical maps (figs 3, 4). the old forms of the names are typically mentioned under the heading ‘history’. place name place name gro#3 western nuussuaq fig. 65 grønne ejland southern disko bugt figs 2, 10 gule ryg eastern disko fig. 124 i igdlokungoak (now illokunnguaq) northern disko fig. 3 igdlokunguak (now illokunnguaq) northern disko fig. 4 igdlunguaq (now illunnguaq) ikorfat northern nuussuaq figs 2, 22 ikorfat fault zone (ik) northern nuussuaq fig. 10 illukunnguaq northern disko fig. 124 illunnguaq southern disko fig. 124 ilugissoq central nuussuaq fig. 2 ilulissat eastern disko bugt figs 2, 10 innanguit southern disko fig. 124 inngigissoq northern disko fig. 124 innaarsuit (also skansen) southern disko fig. 124 ippigaarsukkløften southern nuussuaq fig. 40 itilli western nuussuaq figs 2, 65 itivnera central nuussuaq fig. 82 itsaku svartenhuk halvø fig. 73 ivissussat southern nuussuaq fig. 40 ivissussat qaqqaat southern nuussuaq figs 6, 40 ivisaannguit southern nuussuaq figs 6, 40 ivnanguit (now innanguit) southern disko fig. 4 j.p.j. ravn kløft northern nuussuaq fig. 22 k kaersuarsuk (now qaarsut) kaersut (now qaarsut) northern nuussuaq fig. 4 kangersooq central nuussuaq fig. 82 kangersôq (now kangersooq) kangilia northern nuussuaq fig. 74 kardlok (now qallunguaq) southern nuussuaq fig. 4 kardlunguaq (now qallunguaq) karsoq (now qaarsut) killuusat southern disko fig. 124 kingigtok (now kingittoq) southern nuussuaq fig. 4 kingigtoq (now kingittoq) kingittoq southern nuussuaq fig. 40 kitdlusat (now killusat) southern disko fig. 4 kome (now kuuk) northern nuussuaq fig. 3 kook (now kuuk) northern nuussuaq fig. 4 kook angnertunek (now annertuneq) northern nuussuaq fig. 4 kudliset (now qullissat) northern disko fig. 3 kûk (now kuuk) northern nuussuaq fig. 21 kûk quamassoq (now kuuk qaamasoq) kussinerujuk northern disko fig. 2 kugannguaq valley northern disko fig. 2 kuugannguaq–qunnilik fault (kq) western nuussuaq fig. 10 kuuk northern nuussuaq fig. 21 kuuk qaamasoq southern disko fig. 124 kuussuaq central nuussuaq fig. 65 kvandalen eastern disko fig. 124 l labrador sea fig. 1 laksedalen eastern disko fig. 124 m majorallattarfik northern nuussuaq fig. 21 maligaat north-west of disko fig. 2 marraat southern disko fig. 124 marraat-1 western nuussuaq fig. 65 marrait (now marraat) marrak (now marraat) southern disko fig. 4 melville bay fig. 1 n naajaat southern nuussuaq fig. 40 naassat central nuussuaq fig. 2 nalluarissat central nuussuaq fig. 82 naujat (now naajaat) southern nuussuaq fig. 4 naujât (now naajaat) niakornak (now niaqornat) northern nuussuaq fig. 3 niakornat (now niaqornat) northern nuussuaq fig. 4 niaqornat northern nuussuaq fig. 74 noursoak halfö (now nuussuaq) fig. 3 nûgârssuk (now nuugaarsuk) nûgssuaq (now nuussuaq) nugsuak (now nuussuaq) nugsuaks-halvö (now nuussuaq) fig. 4 nuugaarsuk northern disko fig. 124 nuuk killeq southern nuussuaq fig. 2 nuuk qiterleq southern nuussuaq fig. 2 nuussuaq figs 2, 10 nuussuaq basin fig. 1 o omenak (now uummannaq) fig. 3 170 place name place name p paatuut southern nuussuaq figs 2, 40 paatuutkløften southern nuussuaq fig. 40 patoot (now paatuut) southern nuussuaq fig. 4 pâtût (now paatuut) pautût (now paatuut) peak 1010 m southern nuussuaq figs 6, 40 pingo (now pingu) pingu eastern disko figs 2, 124 pingunnguup kuua western nuussuaq fig. 65 q qaarsut northern nuussuaq figs 2, 22 qaarsutjægerdal central nuussuaq fig. 113 qagdlúnguaq (now qallunnguaq) qallorsuaq (keglen) southern nuussuaq fig. 40 qallunnguaq southern nuussuaq fig. 40 qardlok (now qallunnguaq) qeqertarsuaq (island) figs 2, 10, 73 qeqertarsuaq (town) southern disko figs 2, 10 qilakitsoq central nuussuaq fig. 82 qorlortorsuaq northern disko fig. 2 quikassaap kuua southern nuussuaq fig. 40 quikavsaup kûa (quikassaap kuua) qutdligssat (now qullissat) qullissat northern disko fig. 2 r ritenbenks kulbrud (now qullissat) northern disko figs 3, 4 s sakkak (now saqqaq) southern nuussuaq fig. 3 sarkak (now saqqaq) southern nuussuaq fig. 4 saqqaq southern nuussuaq fig. 2 saqqaqdalen southern nuussuaq figs 2, 40 sarfâgfik northern nuussuaq fig. 21 sarqaq (now saqqaq) scaphitesnæsen central nuussuaq fig. 113 serfat northern nuussuaq fig. 74 sill sø central nuussuaq fig. 113 sinigfik (now siniffik) southern disko fig. 4 sinnifik (now siniffik) southern disko fig. 3 skandsen (now skansen) skansen southern disko figs 2, 124 skarvefjeld southern disko fig. 2 slibestensfjeldet northern nuussuaq fig. 22 sorte hak central disko figs 2, 124 stordal central disko fig. 2 svartenhuk halvø figs 2, 10, 73 t talerua northern nuussuaq fig. 22 tartunaq southern nuussuaq fig. 40 teltbæk fault central nuussuaq fig. 113 tuapassuit northern nuussuaq fig. 21 tuapaat southern disko fig. 124 tuapaat qaqqaat southern disko fig. 124 tunoqqu central nuussuaq fig. 82 tunorqo (now tunoqqu) tunorsuaq northern nuussuaq fig. 74 tupaasat southern nuussuaq figs 2, 6 turritellakløft central nuussuaq fig. 113 u ujaragsugsuk (now ujarassussuk) northern disko fig. 4 ujarattoorsuaq northern nuussuaq fig. 22 ujarasusuk (now ujarassussuk) northern disko fig. 3 ujarasussuk northern disko fig. 124 ukalersalik western nuussuaq fig. 65 umanak (now uummannaq) fig. 4 umiivik-1 svartenhuk halvø fig. 73 umiiviup kangerlua svartenhuk halvø fig. 73 umiussat (now umiusat) umiusat southern nuussuaq fig. 40 upernivik næs fig. 33 upernivik ø figs 2, 10 uppalluk, giesecke monument southern nuussuaq figs 6, 40 uummannaq figs 2, 10 uummannaq fjord fig. 10 v vaigat figs 2, 10 vesterfjeld northern nuussuaq fig. 22 171 place name place name geological survey of denmark and greenland bulletin 20, 2010, 59–62 59 south-east greenland between 62°n and 67°n is one of the lesser known regions in greenland, having seen only limited geological investigations and only few detailed ones, with the skjoldungen alkaline igneous province as a notable exception (nielsen & rosing 1990). systematically collected geoscientific data are scarce; however, such data are essential as a basis for geological models and for evaluation of the mineral potential. in order to open up the region for exploration, the greenland bureau of minerals and petroleum financed a two-year, mainly geochemical programme for 2009 and 2010, which is an initial part of a five to six year project that involves subsequent geophysical surveys, a geological programme and a full-scale resource assessment of the region. the primary objective of the initial geochemical programme is to collect sediment samples for analysis of chemistry and indicator minerals. supplementary to this, surface water for chemistry is collected and radiometric spectra of representative lithologies are measured. geological reconnaissance field work focussing on selected key areas is also carried out. during the past 10–15 years, the geological survey of denmark and greenland (geus) has carried out resource assessments of the palaeoproterozoic orogens and mobile belts in south and central west greenland, and most recently, assessments of the north atlantic craton in west and southern west greenland were conducted. the current assessment of south-east greenland will provide the last major contribution needed for a detailed understanding of the palaeoproterozoic and archaean geological evolution of the entire southern greenland, enabling us to develop new geological models for the region. the 2009 field work focused on the area between timmiarmiit kangertivat and bernstorff isfjord (fig. 1). regional geology the archaean north atlantic craton in south-east greenland is bounded to the south and north, respectively, by the palaeoproterozoic ketilidian and nagssugtoqidian mobile belts. the eastern part of the latter was formerly denoted the ammassalik mobile belt (andrews et al. 1973; escher & nielsen 1983). an overview of the region was provided by chadwick et al. (1989). the mineral resource assessment project, south-east greenland: year one bo møller stensgaard, jochen kolb, troels f.d. nielsen, símun d. olsen, llewellyn pilbeam, diana lieber and anette clausen inland ice st o greenland skjoldungen alkaline province 43°w 63°n 43°w 63°n 41°w ns skjoldungen kattertooq tk timmiarmiit nz sap southern zone 20 km bernstorff isfjord nephelinitic-carbonatitic complex (singertât) mafic granulite, minor paragneiss, metaperidotite, amphibolite grey tonalitic to granodioritic gneiss tonalitic to granodioritic gneiss, locally agmatitic sample gabbro, diorite, monzonite, locally deformed granite and syenite ultramafic rocks, locally deformed syenite, locally deformed granite, diorite, locally deformed archaean metamorphic rocks fig. 1. geological map of part of south-east greenland showing the distribution of sediment samples, surface-water samples, indicator mineral samples and radionetric measurements (black dots). the map is based on escher (1990). nz: northern zone. ns: nordre skjoldungesund. sap: skjoldungen alkaline province. tk: timmiarmiit kangertivat. © geus, 2010. geological survey of denmark and greenland bulletin 20, 59–62. open access: www.geus.dk/publications/bull 6060 the craton is dominated by orthogneisses with subordinate belts and slivers of supracrustal rocks (up to 1 km in width and several tens of kilometres along strike), and with lateto post-tectonic alkaline intrusions in the skjoldungen area. the peak metamorphic facies is granulite grade. retrogression from granulite facies to amphibolite facies assemblages is common. based on lithological variation and regional structural grain the craton is subdivided into a northern zone (nz), a skjoldungen alkaline province (sap) and a southern zone (sz; table 1; fig. 1). the orthogneiss is dominated by an early tonalitic generation intruded by synto late-tectonic tonalite and granodiorite, which have subsequently been intruded by post-tectonic granitic to granodioritic sheets (andrews et al. 1973; escher & nielsen 1983). the early tonalitic gneiss is characterised by an agmatitic fabric with centimetre to metre scale, rounded to angular fragments of amphibolite, meta-diorite and metaultramafic rocks. radiometric age determinations of four gneiss samples (k/ar dating on either biotite or hornblende) yielded cooling ages of 2688–2335 ma (bridgwater 1971). a granulite facies migmatitic gneiss gave a protolith age of 2781 ± 6 ma (zircon u/pb shrimp dating; friend et al. 1996). the supracrustal units comprise amphibolite, metapelite (biotite schist, garnet-quartz gneiss and biotite-garnet-sillimanite-corderite schist), calc-silicate rock and meta-ultramafic rocks including meta-dunite. a 2.8 ga zircon diffusion age is believed to represent the age of metamorphism and hence a minimum age for the formation of the supracrustal rocks (andrews et al. 1973). a sample of an ordinary amphibolite unit yielded a cooling age of 2445 ± 45 ma (biotite k-ar dating, bridgwater 1971). preparation for field work prior to the field work, in 2008 and the first half of 2009, existing data and literature from the region were compiled in digital format and the existing 1:500 000 scale geological map of the region (escher 1990) was modified, updated, and digitised (fig. 1). evenly distributed sampling stations were selected from (1) processed remote sensing data from which drainage systems and catchment basins were defined, (2) satellite images, (3) aerial photographs and (4) topographical data. all information and data are stored in a gis database that will evolve in the coming years and finally comprise all data and observations from the region. regional sampling procedure the regional sampling was carried out by two or three man teams using a helicopter or zodiacs. the chartered vessel m/v fox served as a base. the preferred material for sediment sampling was fine-grained stream sediment from first or second-order streams. in areas without drainage systems, sediment was collected from drift or scree slopes. the rationale behind sampling sediments is that their composition reflects the bedrock as well as results of possible mineralisation processes in the catchment area. the sediment samples were sieved and split in geus laboratories and the fine fraction (<0.1 mm) was analysed for 62 different elements at activation laboratories ltd., canada. sediment samples were collected at 506 locations (fig. 1). in places with suitable drainage systems, the sediment samples were supplemented by surface water samples, with two samples collected at each locality. measurements of ph tonalitic to granodioritic gneiss, locally agmatitic, supracrustal sequences, alkaline intrusive rocks (granitic to syeniticmonzonitic gabbro, carbonatitenephelinite rocks) table 1. differences between three main zones of the archaean craton in south-east greenland and encountered mineralisation subdivision of the craton main lithologies regional structural grain encountered mineralisations and preliminary mineral potential northern zone (nz) numerous layers of supracrustal rocks dominated by mafic granulites. grey tonalitic to granodioritic gneiss n–sto e–w-trending foliation and folds with w-plunging fold axes orthomagmatic ni, cu, pge and au hydrothermal alteration zones with au skjoldungen alkaline province (sap) strong nw–se-trending foliation late conjugate sets of nw–seand sw–netrending foliation alkaline and carbonatite intrusionrelated nb, ree, u and th. hydrothermal alteration zone with sulphides and quartz veins. porphyry cu, sn and mo deposits iron-oxide copper-gold deposits southern zone (sz) agmatitic, tonalitic to granodioritic gneiss few restricted supracrustal sequences n–sto ne–sw-trending foliation and folds with se-plunging fold axes hydrothermal alteration zone with sulphides and quartz veins. alkaline intrusion-related nb, ree, u, th 61 and conductivity were carried out on one of the samples in the base camp. the other sample, which is used for geochemical analysis of 64 elements at the activation laboratories ltd. in canada, was acidified before storage. water samples were collected at 379 localities. a total of 138 coarse-grained sediment samples was collected for indicator mineral analysis from local drift. the rationale behind this type of sampling is that erosion products of a distinct or mineralised rock type can be traced in the drift, and that specific indicator minerals are diagnostic for the specific rock or process. for example, some minerals can be specific for kimberlite that may carry diamonds. material was collected with a spade from below the vegetation, if present. most commonly the drift was till, but locally glaciofluvial sediment was sampled. the material was sieved on a 6.35 mm screen fitted on top of a 20 litres bucket until a sample of c. 5 l was reached. the sieved material was then mixed and transferred to a 5 l container for storage and shipment to the laboratory. the mineral grains will be evaluated by overburden drilling management ltd., canada, for occurrence of kimberlite, au-ni-cu and platinum-group element (pge) indicator minerals. gamma-ray measurements were undertaken at the sample sites, using a portable multichannel gamma-ray spectrometer in order to determine total gamma radiation and concentration variations of k, u and th. finally, representative rock types for geochemistry, petrological investigations and age determinations were collected. preliminary results based on the reconnaissance work, several types of mineralised rock were identified. all encountered mineral occurrences need more work to establish their settings and to assess their potential. small, metre-scale lenses of sulphide-bearing, metamorphosed ultramafic rocks, often hosted within mafic rock sequences (‘supracrustal units’) were found in the northern zone (fig. 1; nz). several rusty horizons, up to tens of metres wide and continuous along strike for tens of kilometres, were seen within mafic to ultramafic rock sequences in both the northern zone and the central skjoldungen alkaline province (fig. 1; nz and sap). the horizons were found to contain disseminated or semi-massive to massive sulphide mineralisation with pyrrhotite, chalcopyrite and pentlandite and were encountered at several places along strike. rock samples from this mineralisation show elevated concentrations of ni, cu and cr (grab samples have yielded up to 0.3% ni, 0.2% cu, 0.5% cr, 24 ppb pt and 162 ppb pd). these rock types may contain occurrences of orthomagmatic ni, cu and pge. quartz veins within hydrothermal alteration zones (pyrrhotite-chalcopyrite-quartz-biotite-garnet) that are 10–20 m wide and traceable for 100 m to several kilometres along strike were found in the skjoldungen alkaline province and in the southern zone. even though samples have only yielded small gold concentrations (maximum 117 ppb au) it is notable that as, a pathfinder element for gold, shows elevated values (up to 1640 ppm as). the alteration zones and the hydrothermal vein systems are potential gold targets. the skjoldungen alkaline province comprises the singertât carbonatite complex (nielsen & rosing 1990; blichert-toft et al. 1995). rock samples from the carbonatite yielded a total rare-earth element (ree) concentration of c. 2500 ppm to41°w 42°w43°w 63°n 63°n 43°w inland ice bernstorff isfjord ns skjoldungen k tk timmiarmiit nz sap southern zone conductivity (μs/cm) 1.7–4.5 4.5–15.6 15.6–31.1 31.1–106.2 ph value 5.9–6.6 6.7–6.9 7.0–7.6 20 km fig. 2. conductivity and ph from surface-water samples. the skjoldungen alkaline province in the central part of the region shows high values of both conductivity and ph. for abbreviations see text to fig. 1. 6262 gether with elevated values of sr, y and ce (up to 2000 ppm sr, 120 ppm y and 1100 ppm ce). the surrounding orthogneisses also show elevated ree concentrations, which may be a result of a hydrothermal halo related to the complex. the syntectonic gabbros, granites and syenites in the skjoldungen alkaline province are rich in magnetite, and the general geological setting may indicate a potential for porphyry copper, tin or molybdenum deposits and iron-oxide– copper–gold deposits. results from regional sampling the field-based gamma-ray measurements of the dominant rock types and the ph and conductivity measurements of surface water provide evidence for significant regional variations. even though the range of both ph and conductivity values is limited, the measurements clearly reflect the intrusive suites of the skjoldungen alkaline province, especially the syenites at kattertooq and western skjoldungen (fig. 2), probably due to outwash of alkaline elements into surface waters from the syenites. this is supported by the gamma-ray measured k content, which is elevated within the skjoldungen alkaline province. elevated ph, conductivity and k content are also seen in the eastern parts of the kattertooq and skjoldungen area, at nordre skjoldungesund and at timmiarmiit in the south. from these data we suggest that alkaline magmatic rocks in the archaean of south-east greenland could be much more widespread than previously recognised. further work the results of the analyses of sediment and surface-water samples and the results from the regional indicator-mineral analysis of material collected in 2009 were released in the spring 2010 (minex 2010). based on these results, the mining industry can evaluate the mineral potential of one of the least explored regions in greenland. the data will also be used by geus and international research partners to improve our understanding of the geology of the region. the programme will continue in 2010 with similar work in the tasiilaq region farther north. many parts of this region are also poorly known and without basal regional data. in addition, aeromagnetic surveys are being considered. the work in 2009 and 2010 provides a basis for more detailed geological work in 2011–2014. the aim of this work is to develop a wellconstrained geological model for the entire region which can justify exploration for ore deposits. acknowledgements we gratefully acknowledge help and support from the skipper niels peter trolle and his excellent crew on m/v fox, from air greenland pilot bertil björk and mechanic benny m. sørensen, and friends in tasiilaq and kulusuk. the greenland bureau of minerals and petroleum supported the project financially. references andrews, j.r., bridgwater, d., gormsen, k., gulson, k., keto, l. & watterson, j. 1973: the precambrian of south-east greenland. in: park, r.g. & tarney, j. (eds): the early precambrian of scotland and related rocks of greenland, 143–156. birmingham university press. blichert-toft, j., rosing, m.t., lesher, c.e. & chauvel, c. 1995: geochemical constraints on the origin of the late archean skjoldungen alkaline igneous province, se greenland. journal of petrology 36, 515–561. bridgwater, d. 1971: routine k/ar age determinations on rocks from greenland carried out for ggu in 1970. rapport grønlands geologiske undersøgelse 35, 52–60. chadwick, b., dawes, p.r., escher, j.c., friend, c.r.l., hall, r.p., kalsbeek, f., neilsen, t.f.d., nutman, a.p., soper, n.j. & vasudev, v.n. 1989: the proterozoic mobile belt in the ammassalik region, south-east greenland, (ammassalik mobile belt): an introduction and re-appraisal. rapport grønlands geologiske undersøgelse 146, 5–12. escher, j.c. 1990: geological map of greenland, 1:500 000, sheet 14, skjoldungen. copenhagen: geological survey of greenland. escher, j.c. & nielsen t.f.d. 1983: archaean gneisses and supracrustal rocks of the tingmiarmiut region, south-east greenland. rapport grønlands geologiske undersøgelse 115, 79–82. friend, c.r.l., nutman, a.p., baadsgaard, h., kinny, p.d. & mcgregor, v.r. 1996: timing of late archaean terrane assembly, crustal thickening and granite emplacement in the nuuk region, southern west greenland. earth and planetary science letters 142, 353–365. minex 2010: new data from south-east greenland. greenland mineral exploration newsletter 36, 2–3. copenhagen: geological survey of denmark and greenland and nuuk: bureau of minerals and petroleum (available at www.geus.dk/minex). nielsen, t.f.d. & rosing, m.t. 1990: the archaean skjoldungen alkaline province, south-east greenland. rapport grønlands geologiske undersøgelse 148, 93–100. authors’ addresses b.m.s., j.k., t.f.d.n., s.o.l. & l.p., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: bmst@geus.dk d.l., weidekampsgade 27, 2. t.v., dk-2300 copenhagen s, denmark. a.c., bureau of minerals and petroleum, p.o. box 930, dk-3900 nuuk, greenland. data article | short jackson et al. 2024: geus bulletin 57. 8339. https://doi.org/10.34194/geusb.v57.8339 1 of 9 a new digital database of ellen louise mertz’s 1924 ‘overview of lateand postglacial elevation changes in denmark’ samuel paul jackson*1 , kristian svennevig1 , kristian kjellerup kjeldsen2 1department for mapping and mineral resources, geological survey of denmark and greenland (geus), copenhagen, denmark; 2department for glaciology and climate, geological survey of denmark and greenland (geus), copenhagen denmark abstract data from an important historic article on lateand postglacial land-level changes in denmark and the accompanying map are presented here in a new digital format. the original data were compiled in 1924 by ellen louise mertz and comprise field observations of the marine limit in denmark made over the late 19th and early 20th centuries. the original tables have been transcribed and expanded into a digital database consisting of 658 entries. the original map sheet has been georeferenced and 392 mapped points have been assigned coordinates. the points are linked to their attributes in the digital data table, making them newly amenable to geospatial analysis in a geographic information system. to demonstrate, we briefly present one such application, namely a reproduction and verification of the isolines of raised beach elevation from the original 1924 map. *correspondence: spj@geus.dk received: 13 dec 2022 revised: 07 nov 2023 accepted: 08 nov 2023 published: 22 mar 2024 keywords: glacial isostatic adjustment, relative sea level, yoldia sea, littorina sea, holocene abbreviations: crs: coordinate reference system gdal: geospatial data abstraction library geus: geological survey of denmark and greenland gis: geographic information system ocr: optical character recognition saga: system for automated geoscientific analyses sdfi: danish agency for data supply and infrastructure tps: thin plate spline utm: universal transverse mercator wms: web map service geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: william colgan (geus, denmark) reviewed by: katie szkornik (keele university, uk) and two anonymous reviewers. funding: see page 8 competing interests: see page 8 additional files: see page 8 introduction in 1924, danish geologist ellen louise mertz compiled a database and map of lateand postglacial land-level changes in denmark (mertz 1924). these were based on evidence of the marine limit, that is, the highest elevations of palaeo-shorelines. her data were extracted from several studies published over the preceding three decades. these studies consist of field observations and measurements made by several danish geologists mapping the country in the late 19th and early 20th century, principally axel jessen, tabular abstract geographical coverage denmark temporal coverage lateand postglacial (latest pleistocene and holocene). original data were collected 1897–1924. subject(s) quaternary geology data format(s) data table of geological observations, including elevation values, with newly extracted coordinates, sourced from an historic study and map, re-tabulated and plotted. data table provided in .csv and .xlsx formats. point and line data provided as geopackage (.gpkg) files with attribute tables. overview map provided in .pdf format. sample collection & analysis digitisation and gis processing of historic data. parameters palaeo-shoreline elevations, strandline elevations, coordinates, deposit and landform types. related publications mertz (1924) and several volumes of the series dgu række i–ii, as listed in database. potential application(s) for these data palaeo-shoreline elevations are one of the constraints on relative uplift and subsidence of the danish landmass over geological time scales. as such they are relevant to studies of glacial isostasy, past sea-level change and as an historical component in calibration of future sea-level change estimates. https://doi.org/10.34194/geusb.v57.8339 https://orcid.org/0000-0002-8825-6160 https://orcid.org/0000-0003-3863-8096 https://orcid.org/0000-0002-8557-5131 mailto:spj@geus.dk https://creativecommons.org/licenses/by/4.0/deed.ast jackson et al. 2024: geus bulletin 57. 8339. https://doi.org/10.34194/geusb.v57.8339 2 of 9 geusbulletin.org vilhelm milthers, kristian rørdam and victor madsen. the entries record the height above sea level of raised beach deposits of various late glacial and postglacial marine phases at several hundred localities across denmark. additionally, a smaller number of measurements of negative elevations from now submerged localities are included. mertz tabulated these measurements and drew a map depicting a large subset of the localities with their respective measurements. on this map, mertz interpolated isolines of elevation for beach deposits of ‘yoldia sea’ age (late glacial) and ‘litorina sea [sic]’ age (postglacial, c. 9500 bp–present), respectively (fig. 1). the original 1924 study, and particularly the set of ‘litorina sea’ isolines, has become a standard reference for lateand postglacial uplift in denmark. google scholar, in 2023, lists 77 citations since 2000, many of which use a version of the mertz (1924) map as part of a figure (see christensen & nielsen 2008; gregersen & voss 2010; hansen et al. 2012; sander et al. 2016). despite still-frequent reproduction and referencing of the isoline map, so far, the original study has only been available as a scan of a physical copy, in which format the data entries are not themselves amenable to computation. in an age of digital analysis, this presents an obstacle to investigation of the data, which this digitisation aims to remove. the 2024 publication centenary is an apt milestone by which to promote the data set and make it more broadly and easily accessible. to this end, we have transcribed the data from mertz (1924) into a digital open access database (jackson et al. 2024). here, we present a newly digitised version of the 1924 map (jackson et al. 2024, file s4), the content of which is distilled in fig. 1. we also demonstrate an example application of modern gis methods to the data set, namely the generation of new computed isolines for comparison with the manual isolines drawn by mertz (fig. 2). study aims, terminology and spelling the intent of this article is simply to present the data contained in mertz (1924) ‘as-is’, albeit, we hope, in a format more amenable to modern use and analysis. we are conscious that, as a century-old article, its content has been built upon, and in some instances superseded, by later advances in the field. this article neither aims to evaluate mertz’s 1924 study nor the views of her contemporaries in the light of this later work and no attempt is made to synthesise the current state of knowledge of these topics. for consistency with the source, we retain the place name spellings used by mertz, even where they have since changed. although the ‘litorina sea’ and ‘zirphæa sea’ are correctly spelled littorina and zirfaea, respectively, we preserve the original spellings of mertz (1924) from here on. danish terms pertaining to the type of deposit in mertz (1924; referred to as ‘aflejringernes beskaffenhed’) are translated to english (table 1). additionally, the concept of ‘tanglinie’ is an important one that warrants brief elaboration. mertz explains that accurate measurements of land uplift since the litorina sea fig. 1 maps of denmark (excluding bornholm) showing isolines of elevation, in metres above sea level, as drawn by mertz (1924) for a: yoldia sea beach deposits (red) and b: litorina sea beach deposits (blue). following the original, solid lines indicate isolines that are based on field data, while dashed lines are inferred. the scanned 1924 original and new digital maps are provided in jackson et al. (2024; files s3 and s4, respectively). na nb https://doi.org/10.34194/geusb.v57.8339 http://www.geusbulletin.org/ jackson et al. 2024: geus bulletin 57. 8339. https://doi.org/10.34194/geusb.v57.8339 3 of 9 geusbulletin.org maximum are obtained by subtracting the height of the uppermost (‘øverste’) tanglinie from the elevation of raised beach deposits. the tanglinie is defined (mertz 1924, p. 5) as the line of highest wave action. this level is typically marked on an active beach by a conspicuous line of debris, such as seaweed. we suggest ‘strandline’ to be an appropriate english translation. there can be several such lines on a beach, hence the specification by mertz of the uppermost strandline. following mertz, it is the present-day uppermost strandline, rather than present-day mean sea level, which provides the most appropriate comparison level to former (now raised) beach deposits, when estimating land uplift. this is because raised beach deposits are a better proxy for the former high-wave action-level than they are for the mean sea level at their time of deposition. the new database includes fields from mertz’s original tables and several new fields, all of which are explained in table 2. historic data and mapping the original data are found in the now discontinued titles published by the geological survey of denmark (danmarks geologiske undersøgelse i–v. række; www. geusjournals.org) between 1897 and mertz’s publication in 1924 (see references in jackson et al. 2024, files s1 and s2). many of the volumes consist of map descriptions covering parts of denmark and include folded maps. data tables are also often included. the volumes describe a wide range of landforms and geological features, of which beach-ridge systems are one part. in compiling her data table, mertz evidently extracted data from both the original tables and from the longform text of the sources. it seems likely that in most, if not all cases, the authors were reporting measurements they themselves had made. in the absence of ‘methods’ sections or equivalent descriptions of exact methods in the texts, however, we are not able to quantify the uncertainties involved, and we therefore limit ourselves to presenting the measurements as they are compiled in mertz (1924). data extraction from the original tables table entries were read from page scans of mertz (1924) and initially written into a microsoft excel (.xlsx) spreadsheet. adobe optical character recognition (ocr) facilitated the copying of text data (locality names and comments), with manual corrections as necessary. numerical data were entered manually but drag-copied where repeated over several rows. danish text is preserved as it appears in mertz (1924), including archaisms and older locality names. fig. 2 maps with new gis-generated isolines interpolated from mertz’s (1924) litorina sea raised beach-elevation point data. the full map is provided in jackson et al. (2024; file s4). the isolines in a are interpolated from only 69 strandline-corrected raised beach elevations (blue dots), plus nine submerged stone age settlements that were used as limiting points (blue crosses). the overall shape and zero-line position are in close agreement with the mertz original, but isolines are more widely spaced towards the north. the alternative map in b shows lines interpolated between all 280 litorina sea raised beach-elevation values mapped by mertz, here ignoring strandline correction, plus nine stone age settlements as before. this interpolation of the uncorrected data reproduces some features of the mertz isolines more closely than the corrected data, notably their convergence over north sjælland and westward curvature over north-west jylland. na nb https://doi.org/10.34194/geusb.v57.8339 http://www.geusbulletin.org/ http://www.geusjournals.org http://www.geusjournals.org jackson et al. 2024: geus bulletin 57. 8339. https://doi.org/10.34194/geusb.v57.8339 4 of 9 geusbulletin.org table 1 summary table providing an overview of the database. id1 mertz source table no. of entries in database2 no. of mapped entries3 descriptive categories4 original danish term suggested english translation n 2501–2623 yoldia sea 97 63 erosionsterrasse wave-cut terrace 46 terrassehak wave-cut notch 16 strandvold beach ridge 11 accumulationsterrasse washover terrace 8 strandgrus beach gravel 7 marine grænse marine limit 2 grusrevle gravel shoal 1 unspecified 6 2701–2711 zirphæa sea 11 9 skallag shell bed 5 kystlinie shoreline 4 erosionsterrasse wave-cut terrace 1 kystskrænt coastal scarp 1 2801–3301 ancylus lake 41 24 tørv peat 37 stubbe og stammer relict tree stumps 4 0101–2409 litorina sea 493 280 strandvold beach ridge 284 havstok beach face 40 strandgrus beach gravel 29 terrassehak wave-cut notch 26 terrasse terrace 20 marine grænse marine limit 20 erosionsterrasse wave-cut terrace 16 kystlinie shoreline 14 strandvold og havstok beach ridge and beach face 10 kystskrænt coastal scarp 8 skallag shell bed 8 grusrevle gravel shoal 6 tørv peat 5 accumulationsterrasse washover terrace 4 unspecified 3 0001–0016 submerged stone age settlements 16 16 total: 658 392 field names5 origin locality dep_abbr find_type note_1924 dgu_ref carried over from original database region division deposit dep_eng elev_min elev_max elev_range e_is_appx sl_min sl_max sl_range sl_is_appx dgu_vol dgu_page expanded from original database table id map_point x_coord y_coord note_2024 elev_map e_minus_sl mertz_page newly assigned or calculated 1id number ranges of original entries in the new data set. 2tally of entries from respective data tables in mertz (1924). 3subset of entries represented by points on the original map. 4breakdown of the descriptive categories assigned to each entry in the original paper, generally a landform or facies. 5breakdown of fields in the data set by origin. the first five fields (locality, dep_abbr, find_type, note_1924, dgu_ref) correspond directly to columns in the original tables, whilst all others either break down the data from a single column over several new columns, or provide information not presented in the original tables. from the original map gis work was carried out in qgis (version 3.22.0; qgis development team 2021). a digital scan of the mertz 1:1  000  000 paper map was georeferenced against web map service (wms) layers published by the danish agency for data supply and infrastructure (sdfi). georeferencing and all subsequent analysis was performed in the universal transverse mercator (utm) 32n (epsg:25832) coordinate reference system (crs) as used by sdfi. georeferencing features were identified from the content of the base map used by mertz (issued originally by generalstabens topografiske afdeling – general staff topographic department 1925). both artificial features (e.g. railway bridges) and natural features (e.g. river mouths) were used. a third-order polynomial transformation was necessary to achieve a satisfactory fit of the paper map to the modern reference layers. a total of 100 control points were used for georeferencing. three reference layers were used: https://doi.org/10.34194/geusb.v57.8339 http://www.geusbulletin.org/ jackson et al. 2024: geus bulletin 57. 8339. https://doi.org/10.34194/geusb.v57.8339 5 of 9 geusbulletin.org 1. the most recent topographic map layer for denmark (danmarks topografiske kortværk 1:25 000), used for orientation and identification of features by place name and geometry (sdfi 2022a). 2. the most recent orthophoto layer for denmark (ortofoto forår, geodanmark), for precise placement of control points (sdfi 2022b). 3. a historic topographic map layer for denmark (lave målbordsblade 1901–1971), to check, where necessary, that control point features have been stable over the past century (sdfi 2022c). the geographical coordinates of the locations on the map are based on the scanned, georeferenced map, with a pixel size of approximately 125 m. we estimate the precision of the mapped points to be at best about an order of magnitude coarser, given that each point is about 8–10 pixels wide, or roughly 1 km at scale. further uncertainty arises from the fact that the placement of points on the original map was based on textual descriptions of their location from the source references. these consist, in some cases, of a cardinal bearing from a named feature (e.g. sw of landmark ‘x’), in other cases of a bearing plus an approximate distance (e.g. 1000 m sw of landmark ‘x’), and in yet others of a simple indication of proximity (e.g. by landmark ‘x’). the size of this uncertainty depends on both the level of detail in the textual description and the level of detail on the base map. this is not straightforward to quantify, but is plausibly at least 1–3 km. data processing generation of new table tabular data were mostly handled in microsoft excel. the data in mertz (1924) are presented in five separate tables, each containing records pertaining to a different phase in the lateand postglacial timeline of denmark. the new database (files s1 and s2, jackson et al. 2024) can be viewed either as a single combined table or filtered on the ‘table’ column for separate viewing of each table. each entry was assigned a unique four-digit id, in which the first two digits identify the geographic subdivision of a given table and table 2 data fields as they appear in the new digital database, alongside an explanation of their meaning. field name description table specifies source table in mertz (1924): yoldia, zirphæa, ancylus, litorina, settlements id unique identifier. the first two digits identify the table and region. region broadest locality specification division intermediate locality specification locality narrowest locality specification map_point number to indicate that the entry is represented with a point on the original map. if not, this cell is empty. numbers are sequential per table, but not unique overall x_coord new value extracted from digitised point geometries y_coord new value extracted from digitised point geometries dep_abbr abbreviation for deposit type as recorded by mertz deposit danish text transcribed from ‘aflejringernes beskaffenhed’ (‘nature of the deposit’). descriptive categories for facies or landforms recorded at the given elevation. dep_eng suggested english translation of descriptive categories (see also table 1). find_type applies to archaeological finds only note_1924 ‘andre oplysninger’ (other information) field transcribed from mertz note_2024 authors’ notes elev_min minimum elevation (højde over havet) value if original table gives a range. otherwise, this cell is empty. elev_max maximum elevation value if original table gives a range. otherwise, the sole value. elev_range the maximum minus the minimum elevation value. e_is_appx binary field, where ‘1’ indicates that the elevation value for the record is marked as approximate (c.) in the original table. otherwise, this cell is empty. elev_map the elevation value as recorded on the original map. in most cases the same as elev_max. sl_min minimum strandline (tanglinie) value if original table gives a range. otherwise, this cell is empty. sl_max maximum strandline value if original table gives a range. otherwise, the sole value. sl_range the maximum minus the minimum strandline value sl_is_appx binary field, where ‘1’ indicates that the elevation value for the record is marked as approximate (c.) in the original table. otherwise, this cell is empty. e_minus_sl new, calculated field for raised beach elevation minus strandline. only applies to cases where both values are available. otherwise, this cell is blank. the new isolines presented in this article are based on these values. mertz_page page number of entry in original mertz table dgu_ref original source reference as listed by mertz dgu_ser original reference series dgu_vol original reference volume dgu_page original reference page https://doi.org/10.34194/geusb.v57.8339 http://www.geusbulletin.org/ jackson et al. 2024: geus bulletin 57. 8339. https://doi.org/10.34194/geusb.v57.8339 6 of 9 geusbulletin.org the last two identify the specific entry within the same. some of the fields from mertz’s tables were expanded into more than one field in the new table to ensure consistent information hierarchy and allow filtering of the data. for example, in mertz’s original table, the ‘locality’ field contains nested sub-headings for regions, regional subdivisions and individual localities, as well as asterisks to indicate which localities are plotted on the map. these were each allocated their own field in the new table (jackson et al. 2024). similarly, mertz’s field ‘publiceret i d.g.u. skrifter’ (meaning ‘published in geological survey of denmark [dgu] series’) lists reference information that was supplemented in the new database with separate fields for series, volume and page number, to make filtering easier. on the original mertz map (file s3, jackson et al. 2024) all points have only a single value shown adjacent to them, even where the table includes a range of values for a given entry. a new field was added for the elevation value as it is displayed next to the points on the original map, which in most cases corresponds to either the table value or the maximum value of a range, but in a few cases corresponds to the minimum value. the determination as to which value is appropriate for the map was made by mertz and is simply carried over. mertz’s ‘højde over havet’ (height above sea level) and ‘nutidens øverste tanglinie’ (uppermost present-day strandline) fields have been expanded to several fields to provide minimum, maximum and range values where relevant. a further new field gives the calculated difference between the height above sea level and present-day strandline at those localities for which both values are provided in the original table. this field supplies the elevation values used to produce a new gis version of the isolines (fig. 2). the elevation values used by mertz in drawing lines were corrected with respect to strandline elevation, as described by mertz (1924, pp. 16–18). our method of isoline contouring differs slightly in that it uses, where possible, the actual georeferenced coordinates and strandline-corrected elevation values of points on the map, rather than the more generally defined areas of the original article. two comment fields are included: the first transcribes, without alteration, the text from mertz’s ‘andre oplysninger’ (other information) field; the second lists the current authors’ notes and is mainly used to acknowledge minor apparent ambiguities in the source tables. new fields were added for the x and y coordinates of the map points (in m, utm zone 32n), obtained by georeferencing of the mertz map as described in the next section. generation of the new digital map localities marked on the mertz (1924) map were digitised as points placed at the visually estimated centre of the dots on the original map, generally while zoomed to between 1:100  000 and 1:10  000 scale. they were saved as a points layer in geopackage (.gpkg) format. to facilitate matching the points to the corresponding table records, they were digitised in the same order as they appear on the original tables. for the litorina table, the points were too numerous for this to be manageable in one sitting, so the map was first sectored manually into smaller geographical areas which, for convenience, followed the pre-existing subdivisions of mertz’s ‘locality’ field. once all points had been mapped, a table join was performed, linking point geometries to all other data fields. x and y coordinates were extracted and added to the table. mertz’s original isolines (fig. 1; file s4) were digitised as polylines by manually tracing over the lines on the georeferenced map. mertz’s differentiation between solid, dashed and dot-dashed lines (see p. 40 in mertz 1924) was represented by the addition of a ‘line style’ field to the layer attribute table, and displayed by means of symbology rules. figures 1 and 2 and a new 1:1 000 000 scale gis-produced map are provided in .pdf format in jackson et al. (2024; file s4). generation of new isolines in addition to the original mertz (1924) data points and isolines, we present a new set of isolines for the litorina data set (fig. 2a, .gpkg files provided in jackson et al. 2024). these are based on the original data points and values but, unlike the manually drawn originals, were automatically generated using gis processing tools. they are intended to provide a comparison between gis methods and the interpreted isolines created for the original study. the new isolines were generated using the thin plate spline (tps) method from the saga (system for automated geoscientific analyses) toolset. an advantage of using a spline interpolation is that it produces reasonable results for areas in which divergent elevation values appear within tight clusters (as occurs in this data set), since the spline is not required to pass through individual data points. however, the spline is consequently an approximation at all points, including at data points. a range of interpolation methods are available, and we have not attempted to assess their respective merits. we limit ourselves to showing that it is indeed feasible to produce isolines that are remarkably congruent with the original using standard gis tools and relatively light parameter tuning. the spline was assigned with a cell size of 1000 m and regularisation value of 1, with all other parameters at default values (processing steps and parameters for replicating the method are provided in the accompanying .txt file in jackson et al. 2024, file s8). the output of the spline tool is a monochrome raster map of elevation values. this was contoured using the contour tool from the geospatial data abstraction library (gdal). the https://doi.org/10.34194/geusb.v57.8339 http://www.geusbulletin.org/ jackson et al. 2024: geus bulletin 57. 8339. https://doi.org/10.34194/geusb.v57.8339 7 of 9 geusbulletin.org contouring interval is 0.5 m, and the range is 0–13 m, as in mertz (1924), for easy comparison. as in the original, the litorina isolines are based on a subset of points whose elevation is corrected by subtracting strandline elevation values. in mertz (1924, p. 17–18) this is based on field measurements from at least 43 localities. the new litorina isolines are based on 69 values calculated by subtracting the strandline from elevation values in the newly digitised data table. also contributing to the interpolation are nine negative elevation values for submerged stone age settlements in the south and west of the map area, which helped to constrain the position of the 0 m elevation isoline. an alternative set of litorina isolines was interpolated using all 280 mapped litorina elevation values, at the expense of correcting for strandline elevation (fig. 2b). data description and main features the data set consists of a list of observations pertaining to lateand postglacial marine limits and transgressed ancient settlements. the key parameter in the database is elevation. most entries record raised beach deposits, which are taken to indicate vertical displacement of a locality since the time of deposition. maximum elevation values in such cases record the highest observed beach deposit at a given locality for a particular marine phase but should be considered asynchronous within each phase. the database (files s1, s2) contains 658 entries, of which 392 are shown on the map produced by mertz (1924). the selection of points represented on the map follows mertz’s own and the placement of points by mertz is presumed to have been interpreted from the text of the sources. the coordinates listed in the new database do not originate from these sources, as they do not provide numerical coordinates, but are extracted from the digitisation of the georeferenced paper map. we have not attempted to interpret the location of further points from the text. the mapped points are symbolised in both the original (file s3) and new map (file s4), according to their source table in mertz (1924). the unmapped points in the tables have no associated coordinates, though their approximate position can generally be inferred from the textual descriptions. here, we provide a brief account of the distribution of each set of mapped points and isolines as shown on both the original and new maps (files s3 and s4). yoldia sea late glacial. shown as red dots. mapped points (63 in total) of yoldia sea deposits occur exclusively in northern jylland – mostly in vendsyssel and around limfjorden, with one point on anholt – as well as on bornholm. their elevation values range from 6.0 m to 57.0 m and generally increase towards the north and east. zirphæa sea after yoldia sea maximum. shown as green dots. all mapped zirphæa deposits (9 in total) are located in a small part of northern vendsyssel, mostly near the northern coast and all within 40 km of each other. elevation values range from 6.0 m to 24.0 m and are highest in the south. ancylus lake timing is the subject of discussion but ending c. 8000 bp (before present). shown as brown dots. the points from the ancylus lake time (24 in total) represent peat deposits and have negative values. they occur in a broad band trending wsw–ene across southern denmark and sjælland, as well as a cluster around limfjorden and one locality on bornholm. they range from –0.3 m (samsø) to –20.0 m (northern germany). litorina sea postglacial, after c. 9500 bp. shown as blue dots. the litorina sea points are by far the most numerous (280 in total). they are also the most geographically widespread, excluding the west and far south of the map area, from which they are absent. their values range from 1.5 m to 14.9 m, generally increasing towards the north and east. some areas are particularly densely sampled, such as around limfjorden and roskilde fjord, and especially the island of samsø. among those points that have strandline values in the database (i.e. those which form the basis of the isolines in fig. 2), limfjorden and roskilde fjord are conversely underrepresented, while samsø remains overrepresented. submerged stone age settlements shown as blue crosses. former settlements, at present submerged, are mapped across the southern part of the map area. of these, nine are ascribed (negative) values, ranging from –1.1 m to –9.0 m, generally becoming more negative towards the south and east. these sites served as limiting points in the automated isoline interpolation. seven localities are mapped that lack an elevation value, though implicitly they have an upper limit of 0 m. isolines yoldia sea isolines (fig. 1) interpreted by mertz (1924) trend wnw–ese across the eastern coast of vendsyssel. they are generally evenly spaced at about 5–7 km, except for the 5.0 m (lowest elevation) isoline, which is https://doi.org/10.34194/geusb.v57.8339 http://www.geusbulletin.org/ jackson et al. 2024: geus bulletin 57. 8339. https://doi.org/10.34194/geusb.v57.8339 8 of 9 geusbulletin.org over 20 km south of the 10 m isoline. isolines on bornholm have the same trend but are roughly twice as closely spaced. litorina sea isolines (fig. 2) interpreted by mertz (1924) trend broadly nw–se, with a gently sinuous course that inflects over the southern kattegat. the 0 m isoline descends from the west coast of jylland to bisect fyn and exits denmark at the island of falster. isolines from 0 m to 5.0 m are roughly evenly spaced at about 10–12 km. isolines above 5.0 m undergo an apparent step change and are spaced more closely, at about 4–6 km, slightly tightening progressively ne. isolines above 5.0 m converge markedly over the northernmost part of sjælland. litorina isolines on bornholm trend wnw–ese and have similarly close spacing to those on northern sjælland. the trend of the litorina isolines in vendsyssel is rotated slightly clockwise relative to that of the yoldia isolines, while the inverse is the case on the island of bornholm. note that isolines are not traced continuously from the danish mainland to bornholm in mertz (1924) or this article. remarks on the new gis-produced isolines computationally interpolating new isolines from the original data can help to assess how well the original manual isolines represent the underlying values. for example, over the island of samsø, in the centre of the map area, there is a pronounced deviation in both the original 1924 (fig. 1) and newly rendered litorina isolines (fig. 2). the appearance of this pattern in the computed isolines can be taken as corroboration that the shape of the manual isolines is indeed reflective of the elevations in this area. conversely, the isolines crossing north sjælland are drawn on the original map as converging strongly eastwards. this convergence appears weaker in the automatically generated isolines based on 69 strandline-corrected elevations (fig. 2a), suggesting that the underlying values do not support such strong convergence. however, when interpolating elevation using all 280 data points, without considering strandline correction, this convergence reappears quite strongly, while also inaccurately offsetting all isolines towards the sw (fig. 2b). this suggests that both corrected and uncorrected values were in fact considered in the construction of the original isolines, despite their being explicitly based only on corrected values. in general, point density appears to influence the shape of isolines, in effect attracting them to converge over densely sampled areas (as on the eastern coast of sjælland, fig. 2). we infer that the accuracy of the isolines is affected by an uneven geographic distribution of points, highlighting the potential value of incorporating additional data to achieve more even data distribution. mertz did not connect her isolines across to bornholm from the danish mainland, and nor have we. we note that any automatic isoline generation clearly lacks mertz’s awareness of the overarching geological context, and this might be expected to account for some divergence in the outcome. we speculate that mertz’s litorina isolines are, for sound geological reasons, a non-uniform treatment of the data, and may not be perfectly replicable from the numerical values. however, overall, if not in detail, the results of the automatic interpolation agree remarkably well with mertz’s interpretation. acknowledgements the authors would like to thank their reviewers for their time and useful comments. additional information funding statement this work is supported by the geocenter denmark project ‘sea-level rise and coastal flooding in denmark: past, future, and policy’, with geocenter partners at the geological survey of denmark and greenland, the department of geoscience at aarhus university and the department of geosciences and natural resource management at the university of copenhagen. authors’ contributions ksv and kkk: conceptualisation, supervision, editing. spj: analysis, original draft. competing interests the authors declare no competing interests. additional files thirteen additional files are available at: https://doi.org/10.22008/ fk2/pi4gxi these comprise: files s1 and s2: the main database table in two formats: file s1 (mertz 1924 data table 2024.csv) and file s2 (mertz 1924 data table 2024.xlsx). file s3: a scanned copy of the original 1924 mertz map (mertz 1924 map original.png). file s4: the newly digitised 1:1 000 000 scale map with gis-generated isolines (mertz 1924 map 2024.pdf). file s5: fig. 1 from this study using original mertz isolines. file s6: fig. 2 from this study using newly produced isolines. file collection s7: geopackage files containing gis point and line data (six files: s7a-f). file s8: instructions on isoline processing steps (mertz 1924 isoline processing steps.txt) references christensen, c. & nielsen, a.b. 2008: dating littorina sea shore levels in denmark on the basis of data from a mesolithic coastal settlement on skagens odde, northern jutland. polish geological institute special papers 23, 27–38. gregersen, s. & voss, p. 2010: irregularities in scandinavian postglacial uplift/subsidence in time scales tens, hundreds, thousands of years. journal of geodynamics 50(1), 27–31. https://doi.org/10.1016/j. jog.2009.11.004 https://doi.org/10.34194/geusb.v57.8339 http://www.geusbulletin.org/ https://doi.org/10.22008/fk2/pi4gxi https://doi.org/10.22008/fk2/pi4gxi https://doi.org/10.1016/j.jog.2009.11.004 https://doi.org/10.1016/j.jog.2009.11.004 jackson et al. 2024: geus bulletin 57. 8339. https://doi.org/10.34194/geusb.v57.8339 9 of 9 geusbulletin.org hansen, j.m., aagaard, t. & binderup, m. 2012: absolute sea levels and isostatic changes of the eastern north sea to central baltic region during the last 900 years. boreas 41(2), 180–208. https://doi. org/10.1111/j.1502-3885.2011.00229.x jackson, s.p., svennevig, k. & kjeldsen, k.k. 2024: ellen louise mertz’s 1924 ‘overview of lateand postglacial elevation changes in denmark’ (v.1). geus dataverse. https://doi.org/10.22008/fk2/pi4gxi mertz, e.l. 1924: oversigt over de senog postglaciale niveauforandringer i danmark. danmarks geologiske undersøgelse, ii. række 41, 49 pp. https://doi.org/10.34194/raekke2.v41.6827 qgis development team 2021: qgis geographic information system, (version release date: october 2021) open source geospatial foundation project. https://www.osgeo.org/projects/qgis/ sander, l., hede, m.u., fruergaard, m., nielsen, l., clemmensen, l.b., kroon, a., johannessen, p.n., nielsen, l.h. & pejrup, m. 2016: coastal lagoons and beach ridges as complementary sedimentary archives for the reconstruction of holocene relative sea-level changes. terra nova 28(1), 43–49. https://doi.org/10.1111/ter.12187 sdfi (danish agency for data supply and infrastructure) 2022a: danmarks topografiske kortværk 1:25 000. https://dataforsyningen.dk/ data/956 (accessed december 2022). sdfi (danish agency for data supply and infrastructure) 2022b: ortofoto forår geodanmark. https://dataforsyningen.dk/data/981 (accessed december 2022). sdfi (danish agency for data supply and infrastructure) 2022c: lave målebordsblade. https://dataforsyningen.dk/data/4519 (accessed december 2022). https://doi.org/10.34194/geusb.v57.8339 http://www.geusbulletin.org/ https://doi.org/10.1111/j.1502-3885.2011.00229.x https://doi.org/10.1111/j.1502-3885.2011.00229.x https://doi.org/10.22008/fk2/pi4gxi https://doi.org/10.34194/raekke2.v41.6827 https://www.osgeo.org/projects/qgis/ https://doi.org/10.1111/ter.12187 https://dataforsyningen.dk/data/956 https://dataforsyningen.dk/data/956 https://dataforsyningen.dk/data/981 https://dataforsyningen.dk/data/4519 a new digital database of ellen louise mertz’s 1924 ‘overview of lateand postglacial elevation changes in denmark’ introduction study aims, terminology and spelling historic data and mapping data extraction from the original tables from the original map data processing generation of new table generation of the new digital map generation of new isolines data description and main features yoldia sea zirphæa sea ancylus lake litorina sea submerged stone age settlements isolines remarks on the new gis-produced isolines acknowledgements additional information authors’ contributions competing interests additional files references figures fig. 1 maps of denmark (excluding bornholm) showing isolines of elevation, in metres above sea level, as drawn by mertz (1924) for a: yoldia sea beach deposits (red) and b: litorina sea beach deposits (blue). following the original, solid lines indicate isolines that are based on field data, while dashed lines are inferred. the scanned 1924 original and new digital maps are provided in jackson et al. (2024; files s3 and s4, respectively). fig. 2 maps with new gis-generated isolines interpolated from mertz’s (1924) litorina sea raised beach-elevation point data. the full map is provided in jackson et al. (2024; file s4). the isolines in a are interpolated from only 69 strandline-corrected raised beach elevations (blue dots), plus nine submerged stone age settlements that were used as limiting points (blue crosses). the overall shape and zero-line position are in close agreement with the mertz original, but isolines are more widely spaced towards the north. the alternative map in b shows lines interpolated between all 280 litorina sea raised beach-elevation values mapped by mertz, here ignoring strandline correction, plus nine stone age settlements as before. this interpolation of the uncorrected data reproduces some features of the mertz isolines more closely than the corrected data, notably their convergence over north sjælland and westward curvature over north-west jylland. tables table 1 summary table providing an overview of the database. table 2 data fields as they appear in the new digital database, alongside an explanation of their meaning. geological survey of denmark and greenland bulletin 20, 2010, 87–90 87 small-scale mining is extraction of metals, precious stones, industrial minerals and other commodities using simple technologies. at a worldwide scale, an estimated 100 million people depend on income from small-scale mining (hinton 2006). in tanzania, there are more than half a million active small-scale miners, most of whom extract gold from placer and hard-rock deposits. apart from providing a livelihood for thousands of households, small-scale mining reduces migration from rural to urban areas. however, small-scale mining is associated with a number of negative effects, because mining activities have severe impacts on both the local environment and the miners’ health. most significantly the widespread use of mercury for gold extraction results in polluted environments and serious health hazards for the miners themselves and for the population in the vicinity of smallscale gold mining settlements (bose-o’reilly et al. 2008a, b; 2010; jønsson et al. 2009). large amounts of mercury are transferred to the environment from small-scale gold mining activities in tanzania (taylor et al. 2005). the mercury remains in the environment and constitutes a severe health hazard, also for generations to come. thus, it is of paramount importance to reduce or, even better, stop the release of mercury from small-scale gold mining. a number of alternative methods have been suggested and tested with limited degrees of success (hilson & van der horst 2002). in 2009 geocenter denmark financed a project to test the feasibility of using borax as a replacement for mercury in small-scale gold extraction in tanzania. gold extraction small-scale gold mining of rock deposits in tanzania is done by sinking shafts and digging tunnels along gold reefs. the mined ore is crushed to walnut size manually or using a jaw crusher and ground in metal drums with hard steel balls or rods, the so-called ball mills. the pulverised material is flushed down a water channel, the bottom of which is covered by a piece of cloth. the heavy particles are caught in the cloth and the light particles end as tailings. this is called sluicing. the heavy fraction from the cloth is treated with mercury (fig. 1). the gold particles amalgamate with mercury and can thus be separated from other heavy minerals. the amalgam is placed in a small iron cup over a fire, the mercury evaporates and the gold is left behind. this gold extraction method is not efficient, so the tailings from sluicing and amalgamation are reprocessed up to ten times in order to recover more gold. the amalgamation method is easy to learn and swift. however, as mentioned above, the method causes serious environmental and health problems and is not particularly efficient with respect to gold recovery. borax – an alternative to mercury for gold extraction by small-scale miners: introducing the method in tanzania peter w.u. appel and jesper bosse jønsson fig. 1. about 100 g of mercury are added to about 5 kg of concentrate from the sluicing. © geus, 2010. geological survey of denmark and greenland bulletin 20, 87–90. open access: www.geus.dk/publications/bull 8888 mercury toxicity during amalgamation the metallic mercury evaporates. some of the vapour is inhaled by people working in the vicinity and may over time cause irreparable damage to their brains. the rest of the mercury vapour gradually precipitates on the ground and enters the drainage system, where it is transformed to methylated mercury by bacteria. methylated mercury is water soluble and enters the food chain causing serious damage to humans who are at the top of the food chain. methylated mercury is extremely harmful to the central nervous system, where it causes tremors, difficulty in walking, tunnel vision, psychological problems and eventually death. there is no cure for permanent mercury poisoning (clarkson et al. 2003). unborn babies are especially prone to damage from methylated mercury. if a pregnant mother has mercury in her body, the foetus ‘sucks’ mercury from her. the nervous system of the foetus is much more sensitive to mercury than that of an adult. a mother with even a low concentration of mercury in her body thus has a high risk of giving birth to mentally or physically disabled children (davidson et al. 1998). gold extraction with borax the use of chemical borax, also known as sodium borate, appears to be one of the more viable ways that have been proposed to reduce or stop the use of mercury by small-scale miners (spiegel & veiga 2010). borax is used for cleaning purposes and is therefore commonly available. the reason for using borax in the smelting process of ore material is that borax reduces the melting point of metals and minerals. under normal field circumstances small-scale miners cannot smelt gold, as they cannot create the high temperature required to smelt the ore. by adding borax to their concentrate, however, they can extract and smelt their gold. gold purchasers already use borax to purify gold with a high content of mercury; however, the method has only recently been applied by small-scale miners. in the benguet area of the northern philippines, around 15 000 small-scale gold miners currently mix their gold concentrate with borax, followed by heating and smelting (leoncio na-oy, personal communication 2010). as a consequence, the mercury usage in the area is minimal and the gold recovery rate quite impressive. the process works as follows: the gold ore is crushed, ground and concentrated as in the case of mercury-based extraction. however, the final product needs to have a very high gold concentration, above 90%, for the borax method to work. this requires skill, practise, and not least time. during testing of the borax method in tanzania, two ways of smelting gold were applied. the first one involved using charcoal and a blower, the second one the use of acetylene gas. charcoal and blower. the gold concentrate is mixed with borax and placed in a plastic bag in a small ceramic bowl filled fig. 2. vigorous burning of charcoal is achieved with a hand-powered blower. fig. 3. molten gold in the centre of the glowing clay bowl and charcoal. 89 with charcoal and some borax, with the recommended ratio between gold concentrate and borax being 1 to 3. the charcoal is ignited and vigorous burning is achieved by using a hand-powered blower (fig. 2). after about 30 minutes the metals and minerals in the concentrate begin to smelt and small drops of gold coalesce at the bottom of the bowl (fig. 3) and can be picked up with the tip of a knife. the advantage of the blower is that it is inexpensive; the disadvantage is that it takes up to half an hour to smelt the gold. acetylene gas. the gold concentrate is mixed with borax and placed in a small plastic bag in a ceramic bowl lined with borax. the gold concentrate –borax mixture is melted with the gas flame and after about 10 minutes the gold melts and coalesces (fig. 4). the advantage of using acetylene gas is that the miners get their gold fast; the disadvantage is that acetylene is expensive and that the gas bottles are heavy to transport. what does it take for the miners to accept the borax method? changing a well-established habit is difficult and requires very good reasons. the habit of using mercury for gold extraction is clearly such a case, because the mercury method is easy to learn and carry out. obviously, this makes the introduction of a new and healthier method a challenge, especially because the borax method requires skill and patience. depending on the skills of the person who prepares the concentrates, the borax method may take between half an hour and an hour longer than the mercury method. the immediate advantage of using borax is that it does not harm the environment or the people within or close to the mining sites, in the quantities necessary for gold extraction. in addition, borax is cheaper than mercury and produces purer gold than that produced with mercury (fig. 5). as small-scale miners are often paid according to the gold content in their gold, the gold produced with borax is likely to provide a better price. however, the question still remains whether these advantages are sufficient to make small-scale gold miners swap from mercury to borax. considering the embedded culture of using mercury, a change from mercury to borax extraction is not likely to come easy. an additional incentive for the miners to convert to the borax method may be needed. the geocenter denmark project that was conducted in two small-scale mining communities in tanzania in may 2009 demonstrated that abandoned tailings from small-scale gold mining may contain a very high gold content, with up to 100 ppm of gold. the average gold grade in the ore mined by the small-scale miners is in the order of 3 to 50 ppm. it was a puzzle how the miners could loose so much gold in spite of repeated processing. the techniques used by small-scale miners are not sophisticated and in general it is believed that around 50% of the gold is lost. nevertheless, tailings with up to 100 ppm gold require an explanation. when small-scale miners treat their concentrate with mercury, they do not recover all the mercury which instead ends up in the tailings. miners know that they are incapable of recovering all the gold in one go and therefore reprocess the tailings in the ball mills, where the mercury becomes pounded to an extent where it looses its ability to coalesce; it turns into what may be termed mercury flour. this flour cannot easily be recovered, which also goes for the gold amalgamated with it. as a result, all the mercury–gold flour is lost to the tailing dumps. awareness by the miners that they lose substantial quantities of gold and income as a result of 1 cm fig. 5. gold extracted by borax (left) is pure, whereas gold extracted by mercury (right) contains up to 10% mercury giving it a paler colour. fig. 4. gold smelting in borax with an acetylene burner. 9090 the mercury technique could be what is needed to facilitate a shift from mercury-based to borax-based gold extraction. lessons learned in rural tanzania the borax method was tested and demonstrated in the two small-scale gold-mining communities londoni and itumbi in tanzania. the tests were carried out by a small-scale miner from the philippines, who uses the borax method on a daily basis. in londoni, a demonstration using charcoal and a blower proved successful and produced a gold tablet of 2.4 grams with a high purity. in itumbi, the test was carried out by using acetylene gas. the test was successful, a tablet of 3 grams was produced, and the project demonstrated that the method is feasible for small-scale gold miners. the small-scale miners from itumbi, a permanent and well-established community, seemed more keen to learn about the mercury-free method than those from londoni, which is a gold-rush settlement, where most residents come from elsewhere. after the demonstration in itumbi, the miners expressed their opinion about the method at a small workshop. on the one hand, most miners recognised the potential of using the borax method, as it would significantly reduce mercury usage and improve the purity of their gold. on the other hand, they mentioned the high price of acetylene gas, the fact that borax is not readily available in the region (this is true and ways of making borax available in that part of the country are needed), the longer time required for preparing the concentrate, and the fact that many miners need to process small quantities of gold (e.g. 0.3 gram) on a daily basis to get food on the table, and that the borax method seemed a bit too advanced for such small quantities. finally, they requested additional instruction and demonstrations in which they could take part before they would be prepared to adopt the new method. conclusions the borax method was received with interest in both smallscale gold-mining communities, but the attitude was more positive in the more permanent settlement of itumbi where people are more concerned about the environment. the following points must be taken into account before borax has a chance of replacing mercury: (1) locally produced, inexpensive blowers or acetylene gas must be easily available, as well as access to the necessary expertise, (2) borax must be readily available, (3) a substantial training programme has to be carried out, (4) a link must be established between the small-scale miners and advisers, preferably the local mining authorities, who can guide the miners when technical problems occur and (5) small-scale miners need to understand the link between the borax method and a higher gold recovery rate. acknowledgements the authors would like to thank the small-scale miners who participated in the testing of the borax method. financial assistance for the research was provided by geocenter denmark. references bose-o’reilly, s., lettmeier, b., gothe, r.m., beinhoff, c., siebert, u. & drasch, g. 2008a: mercury as a serious health hazard for children in gold mining areas. environmental research 107, 89–97. bose-o’reilly, s., lettmeier, b., roider, g., siebert, u. & drasch, g. 2008b: mercury in breast milk – a health hazard for infants in gold mining areas? international journal of hygiene and environmental health 211, 615–623. bose-o’reilly, s., drasch, g., beinhoff, c., tesha, a., drasch, k., roider, g., taylor, h., appleton, d. & siebert, u. 2010: health assessment of artisanal gold miners in tanzania. science of the total environment 408, 796–805. clarkson, t.w., magos, l. & myers, g.j. 2003: the toxicology of mercury – current exposures and clinical manifestations. new england journal of medicine 349, 1731–1737. davidson, p.w. et al. 1998: effects of prenatal and postnatal methylmercury exposure from fish consumption on neurodevelopment: outcomes at 66 months of age in the seychelles child development study. journal of american medical association 280, 701–707. hilson, g. & van der horst, r. 2002: technology, managerial, and policy initiatives for improving environmental performance in small-scale gold mining industry. environmental management 30, 764–777. hinton, j.j. 2006: communities and small scale mining: an integrated review for development planning, 413 pp. washington: world bank group. jønsson, j.b., appel, p.w.u. & chibunda, r. 2009: a matter of approach: the retort’s potential to reduce mercury consumption within small-scale gold mining settlements in tanzania. journal of cleaner production 17, 77–86. spiegel, s.j. & veiga, m.m. 2010: international guidelines on mercury management in small-scale gold mining. journal of cleaner production 18, 375–385. taylor, h., appleton, j.d., lister, r., smith, b., chitamweba, d., mkumbo, o., machiwa, j.f., tesha, a.l. & beinhoff, c. 2005: environmental assessment of mercury contamination from the rwamagasa artisanal gold mining centre, geita district, tanzania. the science of the total environment 343, 111–133. authors’ addresses p.w.u.a., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: pa@geus.dk j.b.j., department of geography and geology, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. geological survey of denmark and greenland bulletin 11, 101-114 101 207pb-206pb dating of magnetite, monazite and allanite in the central and northern nagssugtoqidian orogen, west greenland henrik stendal, karsten secher and robert frei pb-isotopic data for magnetite from amphibolites in the nagssugtoqidian orogen, central west greenland, have been used to trace their source characteristics and the timing of metamorphism. analyses of the magnetite define a pb-pb isochron age of 1726 ± 7 ma. the magnetite is metamorphic in origin, and the 1726 ma age is interpreted as a cooling age through the closing temperature of magnetite at ~600°c. some of the amphibolites in this study come from the naternaq supracrustal rocks in the northern nagssugtoqidian orogen, which host the naternaq sulphide deposit and may be part of the nordre strømfjord supracrustal suite, which was deposited at around 1950 ma ago. pb-isotopic signatures of magnetite from the arfersiorfik quartz diorite in the central nagssugtoqidian orogen are compatible with published whole-rock pb-isotopic data from this suite; previous work has shown that it is a product of subduction-related calc-alkaline magmatism between 1920 and 1870 ma. intrusion of pegmatites occurred at around 1800 ma in both the central and the northern parts of the orogen. pegmatite ages have been determined by pb stepwise leaching analyses of allanite and monazite, and source characteristics of pb point to an origin of the pegmatites by melting of the surrounding late archaean and palaeoproterozoic country rocks. hydrothermal activity took place after pegmatite emplacement and continued below the closure temperature of magnetite at 1800– 1650 ma. because of the relatively inert and refractory nature of magnetite, pb-isotopic measurements from this mineral may be of help to understand the metamorphic evolution of geologically complex terrains. keywords: pb isotopes, magnetite, nagssugtoqidian orogen, palaeoproterozoic, pegmatites, pb stepwise leaching, supracrustal rocks ______________________________________________________________________________________________________________________________________________________________________________ h.s. & k.s., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: hst@geus.dk r.f., geological institute, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. as part of the research programme 2000–2003 in the nagssugtoqidian orogen of west greenland by the geological survey of denmark and greenland (geus), an assessment was made of the mineral resource potential of the region between maniitsoq (sukkertoppen; 66°n) and the southern part of nuussuaq (70°15′n; stendal et al. 2004). the present study comprises pb-isotopic analyses of magnetite from amphibolites, hydrothermally altered amphibolites, the arfersiorfik quartz diorite (see below), skarn, ultramafic rocks and pegmatites. magnetite was chosen as a medium for analysis because of its abundance in amphibolites, even though the concentration of pb in magnetite is generally low. in addition, an attempt was made to date monazite and allanite from pegmatites by the pb stepwise leaching (pbsl) technique (frei & kamber 1995). the pb-isotopic study of the amphibolites covers the attu, kangaatsiaq and qasigiannguit regions (fig. 1). the analysed pegmatites are from the nordre © geus, 2006. geological survey of denmark and greenland bulletin 11, 101–114. available at: www.geus.dk/publications/bull 102 strømfjord (nassuttooq), attu and qasigiannguit areas. the aims of the study were (1) to use pb-isotopic signatures of magnetite in an attempt to outline the metamorphic history of the region; (2) to characterise the hydrothermal overprinting in terms of its timing and pb source; and (3) to place the results within the evolutionary frame of the nagssuqtoqidian orogen. regional geological setting the study region comprises the palaeoproterozoic nagssugtoqidian orogen, a major collisional belt situated just north of the north atlantic craton (van gool et al. 2002), as well as the southernmost part of the contemporaneous rinkian fold belt (garde & steenfelt 1999; connelly et al. 2006). most of the region consists of archaean orthogneisses, variably reworked during the nagssugtoqidian and rinkian tectonothermal events. several thin belts of supracrustal rocks occur within the reworked archaean gneiss terrain of the nagssugtoqidian orogen (fig. 1). granitoid rocks and numerous pegmatites intrude the gneisses. formations of palaeoproterozoic age are limited to the sisimiut igneous suite, arfersiorfik quartz diorite, and minor supracrustal sequences including the naternaq supracrustal belt (connelly et al. 2000; thrane & connelly 2006, this volume). the metamorphic grade is amphibolite facies, except for an area south of ataneq in the south-western part of the northern nagssugtoqidian orogen (nno; fig. 1) and in most of the central nagssugtoqidian orogen (cno), where granulite facies rocks predominate. the gneisses are intensely folded and show a general e–w to ne–sw strike. deformation of the archaean gneisses in the nno greenland inland ice greenland canada 51° ussu itnordre strømfjord arfersiorfik aasiaat qasigiannguit kangaatsiaq attu sisimiut kangerlussuaq naternaq n ag ss ug to qi di an o ro ge n sn o c n o nssz itz n n o inland ice 0 50 km tt t t t t ttt disko bugt 68° jakobshavn isfjord kangersuneq ataneq quaternary surficial deposits basalt metasedimentary rocks palaeoproterozoic palaeogene quaternary sisimiut charnockite arfersiorfik quartz diorite archaean, variably reworked granodioritic and granitic gneiss orthogneiss dioritic gneiss metasedimentary rocks amphibolite (including proterozoic components) t t thrust 457785457785 481049481049 481058481058 484883484883 484890484890 446633446633 446632446632 446626446626 481087481087 446603-04446603-04 44662344662320017362001736 485178485178 485179485179485193485193 223736223736 225348225348223746223746 481070481070 fig. 1. geological map of central west greenland, modified from van gool et al. (2002). red dots with numbers refer to samples analysed. cno, central nagssugtoqidian orogen; nno, northern nagssugtoqidian orogen; sno, southern nagssugtoqidian orogen; itz, ikertôq thrust zone; nssz, nordre strømfjord shear zone. 103 decreases gradually northwards, from high-strain to more open structures in the archaean rocks. steeply and shallowly dipping shear and fault zones are common in contact zones between different rock types. major fault and shear zones generally strike nne–ne. the gneisses of the nno are late archaean, with ages between 2870 and 2700 ma (kalsbeek & nutman 1996; connelly & mengel 2000; thrane & connelly 2006, this volume). however, older rocks with ages ~3150 ma appear to be present in the attu area (stendal et al. 2006, this volume). only a few younger palaeoproterozoic ages have been obtained from the nno, including an undeformed pegmatite between attu and aasiaat with an intrusion age of about 1790 ma (connelly & mengel 2000). the geological history of the study area can be summarised as follows (van gool et al. 2002): • deposition of supracrustal rocks: 2200–1950 ma • continental breakup – the kangâmiut dyke swarm: 2040 ma • drifting – sediment deposition (supracrustal rocks) in the nordre strømfjord area: 2000–1920 ma • subduction – calc-alkaline magmatism, giving rise to the sisimiut and arfersiorfik igneous suites: 1920–1870 ma • peak metamorphism during collision (d1 and d2): 1860–1840 ma • large scale folding (d3): ~1825 ma • shearing in steep belts (d4): ~1775 ma • slow cooling following the shearing, with closing temperature of rutile (420°c) at around 1670 ma (connelly et al. 2000). based on 40ar-39ar and u-pb data for several minerals, willigers et al. (2001) estimated cooling temperatures around 500°c at ~1700 ma, 410°c at ~1640 ma, and 200°c at ~1400 ma. previous investigations the geological survey, university research groups as well as exploration companies have been working in central west greenland for decades and have collected significant amounts of data on the mineral potential of the region (stendal et al. 2002, 2004; stendal & schønwandt 2003; schjøth & steenfelt 2004; steenfelt et al. 2004). whole-rock pb-pb, rb-sr and sm-nd isotopic data from the study area have been presented by kalsbeek et al. (1984, 1987, 1988), taylor & kalsbeek (1990) and whitehouse et al. (1998), while e.g. kalsbeek & nutman (1996), connelly & mengel (2000), connelly et al. (2000), hollis et al. (2006, this volume) and thrane & connelly (2006, this volume) have published zircon u-pb geochronological data. pb-isotopic work has been carried out on sulphide separates, mainly pyrite, from a mineralisation in the disko bugt region north of the study area (stendal 1998). in the latter study, two distinct mineralisation types in the archaean rocks were identified – a syngenetic, and at least one epigenetic type of ore formation. pb-isotopic data of sulphides from proterozoic rocks yield a well-defined linear trend in a pb-pb isochron diagram, with a slope corresponding to an age of ~1900 ma, and indicative of a primitive (i.e. low µ) source character of pb in that mineralisation. local geology and descriptions of the investigated rocks during this study pb-isotopic analyses were carried out on magnetite from amphibolite (four samples), banded iron formation (one sample), hydrothermally altered amphibolite and calc-silicate skarn rock (four samples), the arfersiorfik quartz diorite (three samples), magnetite skarn (one sample), ultramafic rock (one sample) and pegmatite (one sample). in addition, one amphibolite, one altered amphibolite and one sample of banded iron formation were subjected to pbsl procedures (frei & kamber 1995), and allanite (two samples) and monazite (three samples) from pegmatites were analysed by pbsl in an attempt to date their emplacement. brief descriptions of the investigated rocks are given below. amphibolitic rocks amphibolites occur together with garnet-mica schists/ gneisses in supracrustal sequences, interlayered with orthogneiss. some amphibolite layers in the gneiss terrain can be followed continuously along strike for up to tens of kilometres. they are heterogeneous in composition. they are found in three associations: (1) rusty weathering, medium-grained garnet amphibolite layers (c. 0.5 m thick) folded together with the orthogneisses, (2) dark, fine-grained amphibolite, occurring as layers up to 10 m thick, and (3) medium-grained, commonly garnetiferous, layered amphibolite. layered amphibolites are the most common, and occur as units up to 200 m thick, although layers only 10–20 m thick are more common. the three different types of amphibolite form separate outcrops and do not occur together. the pb-isotopic analyses reported in this paper refer to magnetite from the layered amphibolites (type 3). 104 the supracrustal sequences consist of garnet-mica schist/gneiss, together with amphibolite (fig. 2a) and rusty weathering layers c. 1 m thick of quartz-garnet rich gneiss with some iron sulphides (1 vol.%). within the layered amphibolite sequences, magnetite-bearing horizons 1–10 m thick occur. the magnetite occurs in laminae 1–10 mm thick, alternating with quartz-feldspar laminae of the same thickness. alteration is common within the layered amphibolites. altered amphibolite some amphibolites have been hydrothermally altered and sulphide mineralised and may contain calc-silicates. this type of amphibolite is dominated by layered garnet-rich amphibolite, interlayered with magnetite-bearing and rusty weathering layers, with disseminated pyrite (fig. 2b). the layers are generally 0.5–2 m thick; in some cases layered amphibolite is intercalated with rusty weathering layers 10–30 cm thick, consisting of quartz-bearing mica schist with iron sulphides and staining of malachite. within the altered amphibolite calc-silicate minerals are found in zones 1–2 m thick or as smaller lenses, comprising hornblende, diopside, garnet and magnetite. a b fig. 2. amphibolite (a) and hydrothermally altered amphibolite (b) from the attu area. 105 banded iron formation at naternaq the supracrustal belt at naternaq (fig. 1) consists of metavolcanic rocks interlayered with pelitic and psammitic schists and gneisses, marble units, exhalites and chert-rich layers with minor quartzite and banded iron formation. in total, these units define a supracrustal sequence up to 3 km thick, which is folded into a major shallowly dipping wsw-trending antiform. the supracrustal sequence can be traced for approximately 30 km along strike and is intruded by granite sheets and pegmatite veins. østergaard et al. (2002) and stendal et al. (2002) give detailed descriptions of the stratigraphy of the supracrustal rocks. the banded iron formation (fig. 3) occurs locally associated with the amphibolite in zones composed of centimetre-thick layers of magnetite and siderite quartz and calcsilicates. the depositional environment is of a sedimentary type comprising true sediments, submarine volcanic rocks and exhalites. a range of variably altered conformable horizons of very fine-grained siliceous and sulphide rich lithologies associated with either amphibolite or marble are interpreted as volcanogenic-exhalitic rocks (østergaard et al. 2002; stendal et al. 2002). arfersiorfik quartz diorite the arfersiorfik quartz diorite (kalsbeek et al. 1987) is located in the eastern part of the fjord arfersiorfik (fig. 1) and covers several hundreds of square kilometres. within the quartz diorite body, magnetite occurs in hornblenderich rocks (hornblende, quartz, feldspar, and chlorite) and often shows paragenetic relation with iron sulphides (predominantly pyrrhotite). the arfersiorfik quartz diorite was emplaced in the period 1920–1870 ma (kalsbeek et al. 1987; connelly et. al. 2000). ultramafic rocks near qasigiannguit an ultramafic body 300 × 300 m large is located on the north side of kangersuneq, forming rusty weathered hills. on its eastern and western sides the ultramafic body is bounded by fault zones, invaded by pegmatites. on its northern and southern sides it is bordered by amphibolite and garnet amphibolite, respectively. because of the penetrative weathering it is difficult to sample fresh material from the ultramafic body. in its centre, an intensely rusty weathered and eroded ‘joint’ zone cuts the ultramafic rocks. this contains 1–10 vol.% magnetite. magnetite-rich skarn at qasigiannguit near qasigiannguit a skarn rock is found in the contact zone striking 66° and dipping 77°se between mica schist and quartzite and a marble-calc-silicate sequence. it comprises magnetite skarn (0.5 m thick) in close contact with the mica schist and quartzite. towards the south-east the magnetite skarn is followed by alternating layers of calcsilicate rocks and marble (including a quartzitic, sulphiderich layer), followed by a pegmatite body. fig. 3. banded iron formation sequence from the naternaq area. 106 a b fig. 4. pegmatites and minerals analysed. a: pink discordant pegmatite and allanite (inset) from the attu area. b: white pegmatite and monazite (inset) from the nordre strømfjord (nassuttooq) area. 107 pegmatites pink pegmatites. throughout the study area, especially in the outer fjord zone from south of attu northward to kangaatsiaq, the country rocks are intruded by granite and by pink pegmatites with alkali feldspar crystals commonly more than 10 cm in size. the pegmatites occur mostly as discordant decimetreto metre-thick bodies within the gneisses, at contacts between major lithological units, and within supracrustal rocks where they are clearly cross-cutting. the dominant minerals in the pink pegmatite are alkali feldspar, quartz, biotite and subordinate allanite, titanite, apatite, magnetite and fe-sulphides (fig. 4a). zonation is occasionally seen with quartz-rich centres bounded by alkali feldspar-rich parts. white pegmatites. white pegmatites are generally concordant (but locally discordant) to the foliation of the adjacent country rocks, typically grey gneiss and supracrustal rocks. the pegmatites are 5–20 m wide and 50–200 m long with a general trend of nw–se all over the nordre strømfjord and ussuit areas. gradational contacts to the host rocks are common. quartz and feldspar dominate the white pegmatites, with garnet, biotite, monazite, magnetite and zircon as characteristic minor constituents. monazite is found as 0.5–5 mm orange crystals that mainly occur in plagioclaseand biotite-rich pegmatites (fig. 4b). monazite crystals are euhedral and occur in lens-shaped layers accompanied by biotite, set in a granoblastic matrix of primarily plagioclase (secher 1980). analytical methods pb isotope analyses for this study were carried out at the danish centre for isotope geology, geological institute, university of copenhagen. mineral fractions were separated from dry split aliquots of crushed and sieved (100– 200 µm) rock powders using a hand magnet, a frantz isodynamic separator and heavy liquid techniques. no further purification was carried out, and the mineral fractions may contain minor proportions of foreign minerals. pb was separated conventionally on 0.5 ml glass columns charged with anion exchange resin, followed by a clean up on 200 µl teflon columns. a standard hbr-hcl solution recipe was applied in both column steps. total procedural blanks for pb amounted to < 120 pg which is considered insignificant for the measured pb-isotopic results, relative to the amount of sample pb estimated from the mass spectrometer signal intensities. isotope analyses were magnetite from amphibolite and banded iron formation 446626 ataneq 68°.061 53°.510 amphibolite 18.815 0.026 15.549 0.022 44.971 0.068 0.973 0.953 446632 ataneq 68°.047 53°.179 amphibolite 28.204 0.042 16.518 0.026 38.467 0.066 0.971 0.906 446623 attu 67°.837 53°.408 amphibolite 17.522 0.025 15.404 0.024 38.344 0.066 0.961 0.880 481070 attu 67°.915 53°.231 amphibolite 17.388 0.019 15.376 0.018 41.241 0.053 0.978 0.936 484883 naternaq 68°.398 51°.941 bif in amphibolite 54.587 0.028 19.316 0.012 36.871 0.030 0.942 0.886 magnetite from altered amphibolite and calc-silicate rock 446633 niaqornaarsuk 68°.217 53°.028 altered amphibolite 26.387 0.505 16.465 0.316 47.423 0.908 0.998 0.999 446603 attu 67°.927 53°.622 altered amphibolite 18.398 0.030 15.620 0.026 41.710 0.074 0.978 0.956 446604 attu 67°.927 53°.622 altered amphibolite 18.256 0.046 15.638 0.042 40.443 0.122 0.930 0.845 484890 naternaq 68°.408 51°.935 calc-silicates (skarn) 16.534 0.009 15.372 0.010 35.562 0.028 0.957 0.913 magnetite from the arfersiorfik quartz diorite 485178 arfersiorfik 67°.970 50°.430 quartz diorite 17.790 0.016 15.469 0.016 36.703 0.049 0.919 0.782 485179 arfersiorfik 67°.967 50°.412 quartz diorite 18.071 0.018 15.498 0.016 36.343 0.041 0.975 0.954 485193 arfersiorfik 67°.956 50°.594 hornblenditic rock 17.007 0.013 15.456 0.013 37.124 0.037 0.959 0.902 magnetite from pegmatite 481087 attu 67°.890 53°.517 pegmatite 28.393 0.043 16.602 0.026 189.024 0.311 0.982 0.971 magnetite from ultramafic rock and skarn 481049 qasigiannguit 68°.801 50°.973 ultramafic rock 25.827 0.039 16.103 0.025 51.334 0.084 0.979 0.949 481058 qasigiannguit 68°.800 51°.169 magnetite skarn 35.124 0.024 17.028 0.013 38.847 0.035 0.968 0.929 table 1. pb isotope ratios of magnetite from different rock types bif: banded iron formation. * errors are 2σ absolute (ludwig 1990). ** r1 = 206pb/204pb versus 207pb/204pb error correlation (ludwig 1990). † r2 = 206pb/204pb versus 208pb/204pb error correlation (ludwig 1990). sample locality latitude longitude rock 206pb ± 2σ* 207pb ± 2σ 208pb ± 2σ r1 ** r2† n w 204pb 204pb 204pb 108 magnetite, banded iron formation within amphibolite 484883, locality 68°.398 n, 51°.941 w [1] 1 n hbr 30' 48.234 0.030 18.615 0.014 37.418 0.032 0.968 0.919 [2] 1 n hbr 1 h 120.002 0.091 26.407 0.022 3.862 0.005 0.978 0.760 [3] 4 n hbr 3 h 94.120 0.363 23.247 0.090 37.841 0.149 0.995 0.987 [4] 8 n hbr 6 h 30.176 0.239 16.809 0.134 35.387 0.282 0.994 0.994 [5] 8 n hbr 12 h 18.799 0.062 15.506 0.052 34.534 0.116 0.991 0.987 [6] hf 12 h 18.799 0.124 15.539 0.103 35.013 0.233 0.995 0.993 magnetite, amphibolite 446632, locality 68°.047 n, 53°.179 w [1] 1 n hbr 30' 19.932 0.026 15.621 0.021 37.462 0.055 0.970 0.920 [2] 1 n hbr 1 h 38.765 0.032 17.639 0.016 39.167 0.040 0.978 0.946 [3] 4 n hbr 3 h 35.439 0.060 17.098 0.030 42.830 0.079 0.982 0.948 [4] 8 n hbr 6 h 37.099 0.150 17.595 0.072 36.972 0.154 0.993 0.979 [5] 8 n hbr 12 h 40.431 0.219 18.100 0.099 34.544 0.189 0.994 0.994 [6] hf 12 h 28.265 0.083 16.910 0.051 34.598 0.106 0.987 0.977 magnetite, altered amphibolite 446633, locality 68°.217 n, 53°.027 w [1] 1 n hbr 30' 25.206 0.021 16.298 0.015 45.932 0.047 0.971 0.914 [2] 1 n hbr 1 h 27.276 0.170 16.492 0.104 47.555 0.300 0.993 0.990 [3] 4 n hbr 3 h 25.359 0.022 16.261 0.016 45.944 0.049 0.975 0.948 [4] 8 n hbr 6 h 26.018 0.024 16.333 0.016 46.646 0.051 0.968 0.936 [5] 8 n hbr 12 h 27.705 0.073 16.539 0.046 48.236 0.138 0.965 0.937 allanite, pegmatite 2001-736, locality 67°.883 n, 53°.523 w [1] 1 n hbr 30' 22.579 0.029 15.860 0.021 137.787 0.193 0.982 0.960 [2] 1 n hbr 1 h 21.528 0.013 15.728 0.011 123.978 0.101 0.971 0.939 [3] 4 n hbr 3 h 79.398 1.929 22.265 0.542 1360.842 33.070 0.999 1.000 [4] 8 n hbr 6 h 2527.779 37.253 295.424 4.391 52501.990 774.512 0.992 0.999 [5] 8 n hbr 12 h 7804.112 133.916 882.216 15.164 161163.864 2767.112 0.999 1.000 [6] hf 12 h 551.839 2.322 74.260 0.322 11002.171 46.823 0.973 0.994 [7] hf 2 d 111.336 0.543 25.175 0.136 1948.881 9.628 0.901 0.991 table 2. pb-pb step leaching data for magnetite, allanite, and monazite in banded iron formation, amphibolite and pegmatites in the nagssugtoqidian orogen code acid time 206pb/204pb ± 2σ * 207pb/204pb ± 2σ * 208pb/204pb ± 2σ * r1 r2 * errors are 2σ absolute (ludwig 1990). for explanations of r1 and r2, see table 1. allanite, pegmatite 457785, locality 68°.834 n, 51°.226 w [1] 1 n hbr 30' 36.910 1.617 17.616 0.772 538.548 23.595 1.000 1.000 [2] 1 n hbr 1 h 44.614 0.389 18.318 0.160 714.245 6.237 0.998 0.999 [3] 4 n hbr 3 h 26.321 0.210 16.341 0.130 270.923 2.164 0.998 0.999 [4] 8 n hbr 6 h 21966.303 527.491 2410.934 57.939 328786.908 7897.797 0.999 1.000 [5] 8 n hbr 12 h 26.193 0.574 17.165 0.390 161.467 3.639 0.964 0.972 [6] hf 12 h 15.058 1.448 15.006 1.443 39.562 3.805 1.000 1.000 [7] hf 2 d 27.000 0.230 15.806 0.207 186.650 2.183 0.651 0.729 monazite, pegmatite 223736, locality 67°.680 n, 52°.565 w [1] 1 n hbr 30' 136.776 2.665 28.301 0.574 6619.883 131.103 0.960 0.984 [2] 1 n hbr 1 h 222.475 3.796 37.330 0.643 12548.515 214.343 0.991 0.999 [3] 4 n hbr 3h 219.227 2.039 36.918 0.346 12871.792 120.474 0.992 0.995 [4] 8 n hbr 6 h 184.897 2.055 33.259 0.373 10581.910 118.035 0.993 0.997 [5] 8 n hbr 12 h 89.680 0.811 23.046 0.213 3554.756 32.275 0.979 0.998 [6] hf 12 h 39.412 0.038 17.328 0.018 88.489 0.099 0.980 0.957 [7] hf 2 d 261.196 1.366 41.737 0.220 175.006 0.931 0.994 0.987 monazite, pegmatite 223746, locality 67°.674 n, 52°.456 w [1] 1 n hbr 30' 20.955 0.097 15.264 0.071 156.248 0.733 0.995 0.994 [2] 1 n hbr 1 h 26.768 0.169 16.158 0.103 315.778 2.010 0.997 0.997 [3] 4 n hbr 3h 215.079 2.560 37.781 0.451 5263.963 62.733 0.997 0.999 [4] 8 n hbr 6 h 2978.834 188.060 343.540 21.694 75533.892 4768.946 1.000 1.000 [5] 8 n hbr 12 h 6392.913 99.502 721.057 11.259 157988.821 2460.811 0.997 1.000 [6] hf 12 h 468.603 5.974 66.525 0.919 10731.171 137.428 0.923 0.996 [7] hf 2 d 2161.547 97.250 254.541 11.458 51062.309 2297.561 1.000 1.000 monazite, pegmatite 225348, locality 67°.833 n, 52°.323 w [1] 1 n hbr 30' 139.662 2.160 29.676 0.459 1656.012 25.630 1.000 1.000 [2] 1 n hbr 1 h 140.039 1.416 29.481 0.299 1544.411 15.642 0.996 0.999 [3] 4 n hbr 3h 877.104 4.121 108.678 0.517 12440.807 58.889 0.990 0.997 [4] 8 n hbr 6 h 25.459 0.418 16.750 0.456 225.146 4.515 0.604 0.820 [5] 8 n hbr 12 h 40549.054 1970.494 4458.008 222.550 602889.998 29313.071 0.973 1.000 [6] hf 12 h 8037.949 261.731 893.317 29.098 122611.734 3993.108 1.000 1.000 [7] hf 2 d 4389.260 60.969 493.058 6.880 66031.253 918.275 0.996 0.999 109 carried out on a vg sector 54-it instrument. fractionation for pb was controlled by repetitive analysis of the nbs 981 standard (values of todt et al. 1993) and amounted to 0.103 ± 0.007% / amu (2 σ; n = 11). stepwise pb leaching (pbsl) experiments followed methods described in frei & kamber (1995). the programmes and parameters of ludwig (1990) were used for the isochron calculations. model first-stage µ1 values were calculated using 4.55 ga for the age of the earth. all age and isotope data in this paper are given with 2 σ precisions. results the pb-isotopic results are given in tables 1–3. the uranogenic pb-isotopic composition of magnetite from the amphibolites (four samples; squares in fig. 5) together with the banded iron formation (naternaq; one sample outside the range of fig. 5) define an isochron with an age of 1726 ± 7 ma (2 σ; mswd = 1.4; model µ1 = 7.89 ± 0.02), which corresponds to a late stage in the metamorphic evolution of the nagssugtoqidian orogen (cf. willigers et al. 2002). this isochron intercepts the stacey & kramers (1975) pb-isotopic growth curve at ~2140 ma. four mineral separates from altered amphibolite, represented by calc-silicate rich phases and by hydrothermally altered and mineralised samples, have pb-isotopic compositions that plot above the 1726 ma isochron (diamonds in fig. 5). this more radiogenic pb-isotopic composition indicates admixture of a more evolved pb component into the alteration fluids. the pb-isotopic compositions of magnetite from an ultramafic rock and a magnetite skarn from the qasigiannguit area plot below the isochron (out5 samples amphibolite/bif magnetite bulk 1726 6.5 1.38 2140 7.89 0.02 484883 bif magnetite pbsl 1756 36 8.70 2401 7.7 0.11 223736 pegmatite monazite pbsl 1797 13 4.44 2925 7 0.05 223746 pegmatite monazite pbsl 1816 16 41.7 2784 7.24 0.12 225348 pegmatite monazite pbsl 1787 11 76.9 2271 7.81 1.00 447783 pegmatite allanite pbsl 1785 9.2 12.9 2335 7.76 0.00 2001-736 pegmatite allanite pbsl 1818 12 53.8 2453 7.66 0.02 bif: banded iron formation; pbsl: pb step leaching. table 3. pb isotope ages, µ1-values, and intercepts with the stacey & kramers (1975) pb-isotopic growth curve for magnetite, allanite, and monazite from amphibolite, bif and pegmatites sample rock mineral method age (ma) ± 2σ mswd intercept (ma) with µ1 ± 2σ stacey & kramers (1975) 1915 16 17 18 20 7 p b/ 20 4 p b 206pb/204pb 15.2 15.3 15.4 15.5 15.6 2000 1600 1200 800 400 0magnetite amphibolite altered amphibolite arfersiorfik this study arfersiorfik kalsbeek et al. (1987) arfersiorfik whitehouse et al. (1998) age = 1726 ± 7 ma mswd = 1.4 fig. 5. 206pb/204pb-207pb/204pb diagram. squares, pb isotope ratios of magnetite from amphibolites; diamonds, magnetite from altered amphibolites (data from sample 446632 outside the range of fig. 5, see table 1). arfersiorfik quartz diorite: circles, this study; crosses, data from kalsbeek et al. 1987; filled triangles, data from whitehouse et al. 1998. the isochron intercepts the stacey & kramers (1975) pbisotopic growth curve (blue) at ~2140 ma. 110 side the range of fig. 5; see table 1), suggesting a slightly more primitive pb source. the uranogenic vs. thorogenic isotopic patterns (not shown in a figure) are complex and do not add to a better understanding of the uranogenic pb-isotopic data. as expected, they reflect differences in u/th ratios among the different samples analysed. the arfersiorfik quartz diorite has been dated at ~1920 ma (kalsbeek et al. 1987). three magnetite samples from this igneous suite have been included in the present study. the uranogenic pb-isotopic compositions of these magnetites (circles in fig. 5) are similar to the whole-rock pbisotopic signatures (crosses in fig. 5; data from kalsbeek et al. 1987). four additional whole-rock analyses (filled triangles in fig. 5; data of whitehouse et al. 1998) show wider scatter than the data of kalsbeek et al. (1987) and the results of this study. the pbsl data obtained on magnetite from three of these samples are shown in fig. 6. a regression for the steps defined by the sample of banded iron formation, 484883 (excluding step 3; table 1) yields a best-fit line with a slope corresponding to an age of 1756 ± 36 ma (mswd = 8.70; model µ1 = 7.70 ± 0.11; lower intercept with the stacey & kramers pb-isotopic growth curve at ~2400 ma), similar to the age obtained from the amphibolites. pbsl analyses of two other samples (446632, amphibolite and 446633, altered amphibolite) are closely scattered around the 1756 correlation line. pbsl data obtained on allanite from a pink pegmatite (sample 2001-736) resulted in a well-defined errorchron with an age of 1818 ± 12 ma (mswd = 53.8; model µ1 = 7.66 ± 0.02; lower intercept with the stacey & kramers pb-isotopic growth curve at ~2450 ma; fig. 7a). the thorogenic vs. uranogenic isotopic pattern (fig. 7b) reveals that essentially only one phase has dominantly contributed pb to the leaching acids, as a nearly perfect linear relationship is indicated by the data points. this points to a more or less constant th/u in the recovered pb fractions. for this reason, the age of 1818 ± 12 ma can be interpreted with great confidence to represent the emplacement age of the pegmatite. the pb-isotopic composition of magnetite (sample 481087, table 1) from this pegmatite plots on the allanite isochron (fig. 7a), indicating preservation of isotopic equilibrium between these two phases. pbsl data on monazite from a white pegmatite (sample 223736) also yield an errorchron, the slope of which corresponds to an age of 1797 ± 13 ma (mswd = 4.44; model µ1 = 7.00 ± 0.05; lower intercept with the stacey & kramers pb-isotopic growth curve at ~2925 ma; fig. 8a). the thorogenic vs. uranogenic isotopic pattern (fig. 8b) again indicates a predominantly single phase that contributed pb to the leaching acids, as the data points define a near perfect linear relationship. consequently, with great confidence, the age of 1797 ± 13 ma is interpreted as the intrusion age of this pegmatite. 12 16 20 24 28 484883 486632 486633 intercept ~ 2401 ma banded iron formation 484883 altered amphibolite 486633 amphibolite 486632 age = 1756 ± 36 ma mswd = 8.70 magnetite step leaching 1200 20 40 1401008060 20 7 p b/ 20 4 p b 206pb/204pb fig. 6. 206pb/204pb-207pb/204pb diagram of step leaching results of magnetite for three samples. the errorchron intercepts the stacey & kramers (1975) pb-isotopic growth curve (blue) at ~2400 ma. 111 three more step-leaching experiments were performed on allanite (1) and monazite (2) separates from other pegmatites (fig. 9). the ages defined by the respective errorchrons are similar to the ones presented above, and are close to 1800 ma. results of the isochron calculations are listed in table 3. discussion the age defined by the pb-isotopic compositions of magnetite from the amphibolites (1726 ± 7 ma) is younger than the latest major tectonometamorphic event in the region (d4, strike-slip shearing and granite intrusion at 1780– 1770 ma; see connelly et al. 2000 and van gool et al. 2002), and may be interpreted as a cooling age after the d4 event. metamorphic conditions in the cno reached temperatures above 650°c at 1800 ma and approximately 540°c by c. 1740 ma (connelly & mengel 2000; connelly et al. 2000; willigers et al. 2001). slow cooling followed with closing temperatures of rutile (420°c) around 1670 ma (connelly et al. 2000). based on 40ar-39ar and u-pb data of several minerals, willigers et al. (2001) estimated cooling temperatures around 500°c at ~1700 ma, 410°c at ~1640 ma and 200°c at ~1400 ma. a continuous magnetite-ulvöspinel solid solution series exists, with exsolution taking place below 600°c (deer et al. 1966; ramdohr 1969). thus, the ages of the magnetite may date 20 7 p b/ 20 4 p b 10 30 50 70 allanite 2001-736 magnetite age = 1812 ± 12 ma mswd = 54 20 8 p b/ 20 4 p b 206pb/204pb 0 4000 2000 8000 10000 0 100 200 300 400 500 600 magnetite a b fig. 7. step leaching 206pb/204pb-207pb/204pb and 208pb/204pb-206pb/ 204pb diagrams of allanite from red pegmatite (sample 2000736). the errorchron intercepts the stacey & kramers (1975) pb-isotopic growth curve (blue) at ~2925 ma. 20 7 p b/ 20 4 p b 10 20 30 40 50 monazite 223736 age = 1797 ± 13 ma mswd = 4.4 20 8 p b/ 20 4 p b 206pb/204pb 0 4000 12000 16000 0 100 200 300 a b fig. 8. step leaching 206pb/204pb-207pb/204pb and 208pb/204pb-206pb/ 204pb diagrams of monazite from white pegmatite (sample 223736). the errorchron intercepts the stacey & kramers (1975) pb-isotopic growth curve (blue) at ~2453 ma. 112 the timing where exsolution in magnetite ceased (< 1800 ma), that is, after peak metamorphic conditions. model first-stage µ 1 values associated with pb-pb isochrons have been used elsewhere in greenland to judge the influence of pb from archaean sources on the pbisotopic characteristics of palaeoproterozoic igneous rocks (e.g. kalsbeek & taylor 1985). rocks derived from proterozoic sources commonly have model µ 1 values around 8, while contamination with pb from archaean sources tends to lower the µ 1 values. the high µ 1 value (7.89; table 3) obtained for the amphibolite isochron and the lower intercept with stacey & kramers (1975) pb-isotopic growth curve at 2140 ma suggest a mainly palaeoproterozoic pb source for the amphibolites. this source is probably also related to the origin of the supracrustal rocks. detrital zircon u-pb ages of metasedimentary rocks of the nordre strømfjord suite (2200–1950 ma; nutman et al. 1999) and the naternaq supracrustal belt (c. 1950– 1900 ma, thrane & connelly 2006, this volume) indicate erosion of a predominantly palaeoproterozoic hinterland. it implies that the stratabound, semi-massive sulphide deposits associated with banded iron formation at naternaq (stendal et al. 2004) were also deposited during palaeoproterozoic time. the results of allanite and monazite pbsl experiments indicate pegmatite formation around 1800 ma in both the cno (at nordre strømfjord) and nno (at attu and qasigiannguit). this is in agreement with ages reported by kalsbeek & nutman (1996) and connelly et al. (2000), which are slightly younger (1780–1770 ma) or within error overlapping those reported here. the pegmatites were emplaced after post-collisional deformation, large scale folding, and shear zone formation (d3) which ended around 1825 ma (van gool et al. 2002). the wide range in model µ 1 values (7.00–7.81) and lower intercepts with the stacey & kramers (1975) pbisotopic growth curve (2271–2925 ma) indicate variable contributions of archaean and palaeoproterozoic country rocks to the petrogenesis of the pegmatites: pegmatite sample 223736 (µ 1 = 7.00; lower intercept at 2925 ma) may largely consist of remelted archaean country rock, whereas sample 225348 (µ 1 = 7.81; lower intercept at 2271 ma) appears to be mainly derived from palaeoproterozoic sources. hydrothermal activity in the region probably continued after the time of pegmatite emplacement and after the magnetite had cooled through its closing temperature (~600°c), which means that the temperatures of the hydrothermal fluids ranged from 650°c to 400°c in the period 1800–1650 ma. the pb-isotopic signatures of the ultramafic rock and 20 7 p b/ 20 4 p b 206pb/204pb 0 100 150 300 200 250 50 0 c monazite 223746 age = 1816 ± 16 ma mswd = 41.7 400 800 1200 1600 2000 2400 2800 20 7 p b/ 20 4 p b 15 16 17 18 19 allanite 457785 age = 1785 ± 9.2 ma mswd = 12.9 206pb/204pb 0 ma 800 1600 a 15 25 35 45 20 8 p b/ 20 4 p b 206pb/204pb 0 200 400 600 800 1000 b 0 2000 4000 6000 8000 monazite 225348 age = 1787 ± 11 ma mswd = 77 fig. 9. step leaching 206pb/204pb-207pb/204pb diagrams of monazite from white pegmatite (samples 225348 and 223736) and allanite from pink pegmatite (sample 447783). the errorchron intercepts of the stacey & kramers (1975) pb-isotopic growth curve (blue) are given. 113 the magnetite skarn from the qasigiannguit area do not lend themselves to deduce whether these formations were formed during the palaeoproterozoic or represent remnants of archaean origin. it has been suggested that many of the epigenetic gold and copper occurrences in the ataa area north-east of disko bugt, about 75 km north of jakobshavn isfjord (fig. 1), are contemporaneous with the peak metamorphism at ~1900 ma in that area (stendal 1998). this 1900 ma metamorphic-hydrothermal event is not reflected in the magnetite pb-isotopic data of the present study area. conclusions pb-isotopic data of magnetite can be related to the general geological evolution of the nagssugtoqidian orogen and are thus a useful tool for studying the metamorphic history of palaeoproterozoic events in west greenland. a drawback of magnetite pb-isotopic analysis, however, is the generally low pb concentration in this mineral, which makes analysis difficult. magnetite in the amphibolites was formed during several stages of metamorphism. the isochron age of ~1726 ma probably represents a cooling age after a prominent late tectonometamorphic event in the region dated at ~1775 ma. the isotopic data suggest a palaeoproterozoic (mantle?) source for the pb in the amphibolites. the nordre strømfjord supracrustal suite, formed by erosion of a similar juvenile palaeoproterozoic hinterland, was deposited between 2000 and 1920 ma. it is suggested that the naternaq sulphide deposit is part of this supracrustal suite. calc-alkaline magmatism related to subduction (1920– 1870 ma; connelly et al. 2000) gave rise to the formation of the arfersiorfik quartz diorite. the pb-isotopic signature of magnetite from these rocks is comparable with that of whole-rock samples. allanite and monazite pbsl analyses yield pegmatite formation ages of ~1800 ma for both the nordre strømfjord, attu and qasigiannguit regions. the formation of pegmatites is therefore post-collisional. the pegmatites were formed by melting of the local country rocks; pbisotopic data indicate that variable proportions of late archaean and palaeoproterozoic age contributed to their petrogenesis. hydrothermal activity continued after pegmatite emplacement and after closure of magnetite, at 1800–1650 ma. acknowledgements the authors acknowledge f. kalsbeek for inspiring suggestions, and p.m. holm and an anonymous reviewer for further improvements of the manuscript. this paper also contains contributions from several colleagues at geus, all of whom are acknowledged for their work. a special thanks goes to k. markussen, attu, for information about the allanite-bearing pegmatite south of the village. references connelly, j.n. & mengel, f.c. 2000: evolution of archean components in the paleoproterozoic nagssugtoqidian orogen, west greenland. geological society of america bulletin 112, 747–763. connelly, j.n., van gool, j.a.m. & mengel, f.c. 2000: temporal evolution of a deeply eroded orogen: the nagssugtoqidian orogen, west greenland. canadian journal of earth sciences 37, 1121–1142. connelly, j.n., thrane, k., krawiec, a.w., & garde, a.a. 2006: linking the palaeoproterozoic nagssugtoqidian and rinkian orogens through the disko bugt region of west greenland. journal of the geological society, london 163, 319–335. deer, w.a., howie, r.a. & zussman, j. 1966: an introduction to the rock-forming minerals, 528 pp. london: longman. frei, r. & kamber, b.s. 1995: single mineral pb-pb dating. earth and planetary science letters 129, 261–268. garde, a.a. & steenfelt, a. 1999: precambrian geology of nuussuaq and the area north-east of disko bugt, west greenland. geology of greenland survey bulletin 181, 6–40. hollis, j.a., keiding, m., stensgaard, b.m., van gool, j.a.m. & garde, a.a. 2006: evolution of neoarchaean supracrustal belts at the northern margin of the north atlantic craton, west greenland. in: garde, a.a. & kalsbeek, f. (eds): precambrian crustal evolution and cretaceous–palaeogene faulting in west greenland. geological survey of denmark and greenland bulletin 11, 9–31 (this volume). kalsbeek, f. & nutman, a.p. 1996: anatomy of the early proterozoic nagssugtoqidian orogen, west greenland, explored by reconnaissance shrimp u-pb zircon dating. geology 24, 515–518. kalsbeek, f. & taylor, p.n. 1985: isotopic and chemical variations in granites across a proterozoic continental margin – the ketilidian mobile belt of south greenland. earth and planetary science letters 73, 65–80. kalsbeek, f., taylor, p.n. & henriksen, n. 1984: age of rocks, structures, and metamorphism in the nagssugtoqidian mobile belt, west greenland – field and pb-isotope evidence. canadian journal of earth sciences 21, 1126–1131. kalsbeek, f., pidgeon, r.t. & taylor, p.n. 1987: nagssugtoqidian mobile belt of west greenland: cryptic 1850 ma suture between two archaean continents – chemical and isotopic evidence. earth and planetary science letters 85, 365–385. kalsbeek, f., taylor, p.n. & pidgeon, r.t. 1988: unreworked archaean basement and proterozoic supracrustal rocks from northeastern disko bugt, west greenland: implications for the nature of proter114 ozoic mobile belts in greenland. canadian journal of earth sciences 25, 773–782. ludwig, k.r. 1990: isoplot for ms-dos – a plotting and regression program for radiogenic isotope data, for ibm-pc compatible computers, version 2.03. united states geological survey, open file report of-88-0557, 40 pp. nutman, a.p., kalsbeek, f., marker, m., van gool, j.a.m. & bridgwater, d. 1999: u-pb zircon ages of kangâmiut dykes and detrital zircons in metasediments in the palaeoproterozoic nagssugtoqidian orogen (west greenland). clues to the pre-collisional history of the orogen. precambrian research 93, 87–104. østergaard, c., garde, a.a., nygaard, j., blomsterberg, j., nielsen, b.m., stendal, h. & thomas, c.w. 2002: the precambrian supracrustal rocks in the naternaq (lersletten) and ikamiut areas, central west greenland. geology of greenland survey bulletin 191, 24– 32. ramdohr, p. 1969: the ore minerals and their intergrowths, 1174 pp. oxford: pergamon press. schjødt, f. & steenfelt, a. 2004: mineral resources of the precambrian shield of central west greenland (66° to 70°15′ n). part 1. compilation of geoscience data. danmarks og grønlands geologiske undersøgelse rapport 2004/16, 45 pp. secher, k. 1980: distribution of radioactive mineralisation in central west greenland. rapport grønlands geologiske undersøgelse 100, 61–65. stacey, j.s. & kramers, j.d. 1975: approximation of terrestrial lead isotope evolution by a two-stage model. earth and planetary science letters 26, 207–221. steenfelt, a., stendal, h., nielsen, b.m. & rasmussen, t.m. 2004: gold in central west greenland – known and prospective occurrences. geological survey of denmark and greenland bulletin 4, 65–68. stendal, h. 1998: contrasting pb isotopes of archaean and palaeoproterozoic sulphide mineralisation, disko bugt, central west greenland. mineralium deposita 33, 255–265. stendal, h. & schønwandt, h.k. 2003: precambrian supracrustal rocks and mineral occurrences, north-east disko bugt. danmarks og grønlands geologiske undersøgelse rapport 2003/24, 57 pp. stendal, h., blomsterberg, j., jensen, s.m., lind, m., madsen, h.b., nielsen, b.m., thorning, l. & østergaard, c. 2002: the mineral resource potential of the nordre strømfjord – qasigiannguit region, southern central west greenland. geology of greenland survey bulletin 191, 39–47. stendal, h., nielsen, b.m., secher, k. & steenfelt, a. 2004: mineral resources of the precambrian shield of central west greenland (66° to 70°15′). part 2. mineral occurrences. danmarks og grønlands geologiske undersøgelse rapport 2004/20, 212 pp. stendal, h., frei, r. & stensgaard, b.m. 2006: a lead isotope study of an archaean gold prospect in the attu region, nagssugtoqidian orogen, west greenland. in: garde, a.a. & kalsbeek, f. (eds): precambrian crustal evolution and cretaceous–palaeogene faulting in west greenland. geological survey of denmark and greenland bulletin 11, 53–60 (this volume). taylor, p.n. & kalsbeek, f. 1990: dating the metamorphism of precambrian marbles: examples from proterozoic mobile belts in greenland. chemical geology 86, 21–28. thrane, k. & connelly, j.n. 2006: zircon geochronology from the kangaatsiaq–qasigiannguit region, the northern part of the 1.9– 1.8 ga nagssugtoqidian orogen, west greenland. in: garde, a.a. & kalsbeek, f. (eds): precambrian crustal evolution and cretaceous– palaeogene faulting in west greenland. geological survey of denmark and greenland bulletin 11, 87–99 (this volume). todt, w., cliff, r.a., hanser, a. & hofmann, a.w. 1993: re-calibration of nbs lead standards using a 202pb + 205pb double spike. terra abstracts 5, supplement 1. van gool, j.a.m., connelly, j.n., marker, m. & mengel, f. 2002: the nagssugtoqidian orogen of west greenland: tectonic evolution and regional correlations from a west greenland perspective. canadian journal of earth sciences 39, 665–686. whitehouse, m.j., kalsbeek, f. & nutman, a.p. 1998: crustal growth and crustal recycling in the nagssugtoqidian orogen of west greenland: constraints from radiogenic isotope systematics and u-pb zircon geochronology. precambrian research 91, 365–381. willigers, b.j.a., krogstad, e.j. & wijbrans, j.r. 2001: comparison of thermochronometers in a slowly cooled granulite terrain: nagssugtoqidian orogen, west greenland. journal of petrology 42, 1729– 1749. __________________________________________________________________________________________________________________________________________________________________________________ manuscript received 28 october 2004; revision accepted 19 december 2005 geological survey of denmark and greenland bulletin 13, 2007, 53-56 53 during the field campaign in the nuuk region, one of the objectives was to describe archaean primary geological environments (hollis et al. 2006). on nunatak 1390, which is part of the tasiusarsuaq terrane (figs 1, 2), a bi modal volcanic succession is preserved and interpreted as former ocean floor. the field investigation included geological mapping and sampling of the volcanic sequence comprising mafic to ul tra mafic rocks, and associated acid volcanic rocks and granite intrusions. the tasiusarsuaq terrane the tasiusarsuaq terrane is dominated by mafic rocks (amphibolite), tonalitic gneiss and granodiorite yielding ages of 2.92– 2.86 ga (fig. 1; schiøtte et al. 1989; friend & nutman 2001; crowley 2002). metamorphic grade ranges from greenschist to granulite facies conditions, with peak metamorphism dated at ~2.79 ga (pidgeon & kalsbeek 1978). the mafic rocks comprise greenschist facies mafic rocks with pillow structures, metagabbroic and ultramafic pods, and dykes. the thicknesses of the mafic to ultramafic sequences vary from 50 m up to more than 1000 m. the tasiusarsuaq terrane rocks are cross-cut by brown-weathering e–w-trending dolerite dykes (up to 30 m wide) with well-developed chil led margins. alterations such as calc-silicate formation are common within the maficultramafic rocks. the pillowed mafic se quences contain intercalations of 1–2 m wide, rusty, sulphide-bearing layers and tourmalinites (exhalites). the sulphides recorded include pyrite, pyrr hotite, chalcopyrite and arsenopyrite. a well-preserved bimodal archaean volcanic succession in the tasiusarsuaq terrane, south-west greenland henrik stendal and anders scherstén © geus, 2007. geological survey of denmark and greenland bulletin 13, 53–56. available at: www.geus.dk/publications/bull fig. 1. geological map of the nuuk region and location of the nunatak 1390 study area (modified from escher & pulvertaft 1995). nunatak 1390 nunatak 1390 is located within the inland ice east of alangorlia and was first described by escher & pidgeon (1976; figs 1, 2). the entire volcanic package (fig. 3; table 1) is north-east-striking and dips steeply to the north-west. although slightly to moderately deformed with lineations, folding, faulting and shearing, the rocks show well-preserved primary textures. stratigraphy the lower mafic pillow sequence shows large deformed pillow structures (50–100 cm across) and pillow breccias with calc-silicate alteration in the matrix between the pillows and in the centre of some pillows. the calc-silicate minerals include epidote, diopside and carbonates and make up to 20 vol.% of the rock. the pillowed sequence is cut by a slightly deformed e–w-trending swarm of mafic dykes (1–5 m thick). the dykes are fineto medium-grained gabbroic or noritic rocks. ultramafic greenstones and soapstones occur between the upper and lower pillow lava sequences. these magnetitebearing rocks were probably originally sills. the upper pillow sequence contains very well-preserved primary structures in pillows, lava flows and ash layers (fig. 4). the least deformed pillow lavas and flows contain relic vesicles. way-up can readily be determined from the pillow structures and consistently youngs to the south. the upper mafic pillow sequence is overlain by a unit of acid volcanic and pyroclastic rocks, including ignimbrites, 80 m in thickness (fig. 5). finegrained, grey to light-coloured porphyritic dykes (0.3–0.8 m wide) cut the volcanic rocks and are interpreted as feeder dykes to the acid rocks. mafic flows and ash layers are intercalated with ultramafic sills, and mafic rusty layers contain sulphides and tourmalinites. the tourmalinite forms an up to one metre thick layer. a prominent hydrothermal zone, strongly silicified and epidotised, strikes parallel with the mafic ash layers. it follows 54 fig. 2. geological map of nunatak 1390. profile a–b is described in table 1 and shown in fig. 3. pd, palaeoproterozoic dyke. for location see fig. 1. a fault lineament, is up to 50 m wide and can be recognised from the light brownish surface colour of the altered rocks. the hydrothermal zone is overlain by a thick sequence (700–800 m) of finely laminated tuff layers. granite intrusions in the tuffs increase in abundance upwards and pass upwards into porphyritic granite with tuff xenoliths. two phases of granite occur: one is porphyritic with kfeldspar phenocrysts up to several centimetres in length; the other is more homogeneous, medium-grained, slightly foliated and muscovite-bearing. parts of the granites, especially in the western part of the exposure, are altered and have a distinct pink coloration due to hematite formation. on a regional scale, it should be noted that western and southern parts of the tasiusarsuaq terrane preserve remnants of volcanic rocks at several localities, probably of similar age to that on nunatak 1390 (escher & myers 1975). geochemistry twenty samples representing most rock types found on nunatak 1390 were analysed by actlabs, canada (research package 4e). altered samples were screened using e.g. k2o/p2o5, and those with anomalous ratios are not considered any further. melanocratic–ultramafic pillow lavas, flows and ash have komatiitic and mgand fe-rich basalt compositions and plot along a well-defined tholeiitic trend (fig. 6a). the acid rock, which intercalates with the tholeiites, forms a loosely defined calc-alkaline group of andesitic to dacitic composition (fig. 6a). the tholeiites are characterised by near chondritic relative ree abundances (mean la/smn = 1.2 ± 0.4 1σ; la/ybn = 1.3 ± 0.3 1σ, n = 9), while the acid rocks are more enriched and varied in their lree (mean la/smn = 5 ± 1 1σ; la/ybn = 24 ± 17 1σ, n = 5; fig. 6b). the different incompatible trace element abundances in the tholeiites and acid rocks persist through all the elements, albeit with more scatter in the mobile elements. important immobile, incompatible element ratios are indicative of potentially different tectonic settings for the tholeiites and the acid rocks. nb/la ratios do not vary with nb or other incompatible elements, which implies insig nif icant effects of fractional crystallisation and insignificant amounts of crustal contamination (fig. 6c). the tholeiite mean nb/la ratio is 0.6 ± 0.1 (1σ; n = 9), while the acid rocks have a substantially lower ratio of 0.14 ± 0.03 (1σ; n = 5). the ratios of the tholeiites are reminiscent of lower crust, while the low ratios of the acid rocks are typical of volcanic arc related rocks (hawkesworth & kemp 2006). even though metamorphic element mobility may disturb e.g. ce/pb ratios, the consistent and low ratios of the acid rocks corro borate an arc origin (fig. 6d), which we postulate for these rocks. the tholeiites are more ambiguous and are akin to morb or island arc tholeiites. the spatial relationship 55 fig. 3. central part of the stratigraphy of profile a–b on nunatak 1390 (see table 1). fig. 4. pillow lava structures in the upper pillow lava sequence. hammer (50 cm) for scale. fig. 5. laminated acid volcanic rock (ignimbrite). inset is an enlargement (inset is about 5 cm wide). between the acid rocks and the tholeiites seems to support a common origin, in which case an arc setting seems most plausible. by analogy with modern arc systems, the two components may reflect input from trench-side (tholeiites) and back arc-side (calc-alkaline) volcanoes respectively, or temporal shifts in the petrogenetic processes. references crowley, j.l. 2002: testing the model of late archean terrane accretion in southern west greenland: a comparison of the timing of geological events across the qarliit nunaat fault, buksefjorden region. precambrian research 116, 57–79. escher, j.c. & myers, j.s. 1975: new evidence concerning the original relationships of early precambrian volcanics and anorthosites in the fiskenæsset region, southern west greenland. rapport grønlands geologiske undersøgelse 75, 72–76. escher, j.c. & pidgeon, r.t. 1976: field mapping of nunatak 1390 m, east of alángordlia, southern west greenland. rapport grønlands geolo giske undersøgelse 80, 84–87. escher, j.c. & pulvertaft, t.c.r. 1995: geological map of greenland, 1:2 500 000. copenhagen: geological survey of greenland. friend, c.r.l. & nutman, a.p. 2001: u-pb zircon study of tectonically bounded blocks of 2940–2840 ma crust with different metamorphic histories, paamiut region, south-west greenland: implications for the tectonic assembly of the north atlantic craton. precambrian research 105, 143–164. hawkesworth, c.j. & kemp, a.i.s. 2006: evolution of the continental crust. nature 443, 811–817. hollis, j.a., schmid, s., stendal, h., van gool, j.a.m. & weng, w.l. 2006: supracrustal belts in godthåbsfjord region, southern west greenland. progress report on 2005 field work: geological mapping, regional hydrothermal alteration and tectonic sections. danmarks og grøn lands geologiske undersøgelse rapport 2006/7, 171 pp. pidgeon, r.t. & kalsbeek, f. 1978: dating of igneous and metamorphic events in the fiskenaesset region of southern west greenland. canadian journal of earth sciences 15, 2021–2025. schiøtte, l., compston, w. & bridgwater, d. 1989: u-pb single-zircon age for the tinissaq gneiss of southern west greenland: a controversy resolved. chemical geology (isotope geoscience section) 79, 21–30. 56 authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: hst@geus.dk fig. 6. a: feot/mgo versus sio2 variation diagram, and jensen cation plot; rhyolite (r), dacite (d), andesite (a), basalt (b), mg tholeiite basalt (m), fe tholeiite basalt (f), komatiitic basalt (kb) and komatiite (k). b: ree distribution of the tholeiitic suite (green) and calc-alkaline suite (red). c: nb versus nb/la of the tholeiitic suite (green) and calc-alkaline suite (red). d: ce/pb ratios variation versus ce of the tholeiitic suite (green) and calc-alkaline suite (red). geological survey of denmark and greenland bulletin 11, 33-52 33 pre-nagssugtoqidian crustal evolution in west greenland: geology, geochemistry and deformation of supracrustal and granitic rocks north-east of kangaatsiaq jean-françois moyen and gordon r. watt the area north-east of kangaatsiaq features polyphase grey orthogneisses, supracrustal rocks and kangaatsiaq granite exposed within a wsw–ene-trending synform. the supracrustal rocks are comprised of garnet-bearing metapelites, layered amphibolites and layered, likewise grey biotite paragneisses. their association and geochemical compositions are consistent with a metamorphosed volcano-sedimentary basin (containing both tholeiitic and calc-alkali lavas) and is similar to other archaean greenstone belts. the kangaatsiaq granite forms a 15 × 3 km flat, subconcordant body of deformed, pink, porphyritic granite occupying the core of the supracrustal synform, and is demonstrably intrusive into the amphibolites. the granite displays a pronounced linear fabric (l or l > s). the post-granite deformation developed under lower amphibolite facies conditions (400 ± 50°c), and is characterised by a regular, ne–sw-trending subhorizontal lineation and an associated irregular foliation, whose poles define a great circle; together they are indicative of highly constrictional strain. the existence of a pre-granite event is attested by early isoclinal folds and a foliation within the amphibolites that is not present in the granite, and by the fact that the granite cuts earlier structures in the supracrustal rocks. this early event, preserved only in quartz-free lithologies, resulted in high-temperature fabrics being developed under upper amphibolite to granulite facies conditions. keywords: archaean, deformation, supracrustal rocks, granite, nagssugtoqidian _______________________________________________________________________________________________________ j.-f.m., department of geolog y, university of stellenbosch, 7602 matieland, south africa. e-mail: jfmoyen@wanadoo.fr g.r.w., marchmyres cottage, breda, alford ab33 8nq, aberdeenshire, u.k. introduction and regional geology the northern part of the nagssugtoqidian orogen (nno) is a domain of predominantly archaean rocks that have been deformed and metamorphosed during nagssugtoqidian orogenic activity at c. 1.8 ga (hollis et al. 2006, this volume; mazur et al. 2006, this volume; thrane & connelly 2006, this volume; van gool & piazolo 2006, this volume). palaeoproterozoic rocks are sparse, and apparently confined to some supracrustal belts, the most prominent one being the naternaq supracrustal belt (østergaard et al. 2002). a few small granitic-pegmatitic plugs and dykes are also related to the palaeoproterozoic evolution. therefore, while the structures probably reflect nagssugtoqidian deformation, the rocks themselves (and their protoliths) largely reflect archaean formations and evolution. among the archaean units, the ubiquitous orthogneissic basement has previously been studied (moyen et al. 2003a; steenfelt et al. 2005); it is mostly made up of classical tonalite-trondhjemite-granodiorite (ttg) gneisses, with minor components either related to ttg partial melting, or to the participation of peridotitic mantle in their petrogenesis. all these components are well known in the archaean, and are probably related to the subduc© geus, 2006. geological survey of denmark and greenland bulletin 11, 33–52. available at: www.geus.dk/publications/bull 34 tion of hot oceanic lithosphere in an arc setting (martin 1986, 1994; moyen et al. 2003b; steenfelt et al. 2005). several components of the gneissic basement have been dated (kalsbeek & nutman 1996; connelly & mengel 2000; thrane & connelly 2002, 2006, this volume), yielding ages in the range 2.9–2.6 ga. supracrustal assemblages are common, and have been mapped in many places in the kangaatsiaq, aasiaat and kangersuneq map sheet areas (marker et al. 1995; mengel et al. 1998; østergaard et al. 2002; van gool et al. 2002a; garde 2004; hollis et al. 2006, this volume). they are of two main types, either amphibolites or metasedimentary rocks, that may be either aluminous, biotite ± muscovite ± sillimanite ± garnet-bearing metapelites, or quartz-rich, psammitic rocks. the age of the supracrustal rocks is, however, poorly constrained. whilst some of them are of demonstrated palaeoproterozoic age (c. 1.95 ga, the naternaq supracrustal belt, østergaard et al. 2002; thrane & connelly 2002, 2006, this volume), others are likely to be of archaean age, for instance anatectic metapelites in saqqarput fjord in the southern part of the kangaatsiaq map sheet area (thrane & connelly 2006, this volume). lastly, small intrusions, plugs and sheets of granite and pegmatite cut across the lithologies described above. some of them have been dated (kalsbeek & nutman 1996; thrane & connelly 2002, 2006, this volume) and yielded late archaean ages (2.7–2.6 ga); it is commonly agreed that most magmatic activity in this region was related to late archaean events, palaeoproterozoic p–t conditions being such that anatexis was hardly achieved in the nno (mazur 2002; piazolo 2002). the very homogeneous and porphyritic kangaatsiaq granite north-east of kangaatsiaq, 15 by 3 km in outcrop size, is among the largest granitic bodies of presumed late archaean age in the southern disko bugt region. altogether, the three components outlined above are representative of the usual trilogy of archaean terranes (windley 1995): grey ttg gneisses; volcanic and volcano-sedimentary deposits (greenstones); and late, k-rich granites. the area east and north-east of the town of kangaatsiaq (fig. 1) is dominated by a synform of supracrustal rocks (mafic and felsic volcanic rocks associated with sediments), into which the kangaatsiaq granite was emplaced. it is, therefore, a good place to study the archaean components and local history in the nno. greenland 68°19' 53°24' kangaatsiaq granite amphibolite layered biotite gneiss of supracrustal origin and aluminous metapelite orthogneiss geological boundaries: established inferred fault c d b a a 81 82 89 92 2 km kangaatsiaq 61 73 52 85 64 58 57 80 75 fig. 1. geological map of the kangaatsiaq granite and surrounding synform, with sample localities from table 1. geology mostly from 2002 field work; some parts are drawn from 2001 data (j.a.m. van gool, g.i. alsop, s. piazolo and s. mazur). a–a, approximate position of section on fig. 2; b, loc. 89, see fig. 3; c, locs 81–82, see fig. 4; d, loc. 80, see fig. 5. 35 previous work previous studies in the kangaatsiaq area included reconnaissance mapping by noe-nygaard & ramberg (1961), 1:250 000 scale mapping by henderson (1969), and visits to key localities during the danish lithosphere centre project (marker et al. 1995; mengel et al. 1998), as a result of which most published ages were obtained (kalsbeek & nutman 1996; connelly & mengel 2000). mapping of the area was predominantly based on coastal exposures, while map information for large parts of the inland areas was based only on photogeological interpretation. therefore, the kangaatsiaq granite, which happens to crop out mostly inland and occupies the high grounds at the core of a synform, was at that time simply considered to be part of the polyphase gneissic basement. the geological survey of denmark and greenland (geus) and its partners undertook more detailed mapping of the kangaatsiaq map sheet in the summer of 2001. this included limited inland work, and the kangaatsiaq granite was recognised for the first time as belonging to the group of the late archaean intrusives. its overall shape was mapped, as well as the supracrustal rocks into which it intrudes. metasedimentary rocks in the area were also sampled, allowing for metamorphic studies (mazur 2002; piazolo 2002). finally, re-evaluation of the area in the summer of 2002 by the present authors led to the refinement of geological boundaries and the production of the map of fig. 1. sampling of the whole supracrustal series was also undertaken. thin sections were cut at université claude-bernard (lyon, france), and samples were analysed at geus using xrf as well as icp-ms (table 1). in addition, other supracrustal rocks from the same area (obtained from a.a. garde, personal communication 2003) have been used for the interpretation presented here, as they show similar geochemical features. map pattern as mentioned in the introduction, the studied area (fig. 1) is mainly made up of three main components: basement orthogneisses discussed by moyen et al. (2003a) and steenfelt et al. (2005), a succession of supracrustal rocks which comprise a sequence of amphibolite and metasedimentary rocks described below (figs 1, 2), and the kangaatsiaq granite, an intrusion of pink, coarse-grained, strongly lineated (l or l > s fabric) granite with k-feldspar phenocrysts. the foliated basement gneisses and the supracrustal rocks, together with early folds and structures, are refolded into a complex synform which is locally overturned, in particular on its north-western rim (see below). the granite occupies the core of the synform; it is intrusive within the top amphibolitic layer of the supracrustal sequence (fig. 3) and is also folded together with the supracrustal rocks. the geometry of the granite suggests that it constitutes a single sheet within the supracrustal unit, and that the original intrusion had an overall flat, laccolithlike shape (fig. 2). we consider that the mapped contact always corresponds to the bottom of the laccolith, and that the top surface has been removed by erosion (fig. 2). nw se kangaatsiaq granite amphibolite 1 km layered biotite gneiss of supracrustal origin with ultramafic layer (schematic) with gabbroic lens (schematic) aluminous metapelites basement orthogneiss amphibolite layers and enclaves in other lithologies (schematic) fig. 2. schematic nnw–sse crosssection across the kangaatsiaq granite and the surrounding synform. the laccolith shape (dashed line) is inferred, see text for details. 36 a b c d photo jfm-2002-5-22 photo jfm-2002-5-22 photo jfm-2002-5-22 photos jfm-2002-5-19–21photos jfm-2002-5-19–21photos jfm-2002-5-19–21 fig. 3. contact of the kangaatsiaq granite and the south-western limb of the synform, loc. 89. the granite clearly intrudes the supracrustal pile, and at the same time occupies the core of the (here, slightly overturned) synform with apparently conformable relationships. a: photomosaic of cliff face, facing east. b: structural interpretation (stippled: pegmatites; rectangle: location of enlargement d). c: lithological interpretation. crosses: granite; dark grey: amphibolite; light grey: layered gneiss. d: detail of a small granitic apophysis which clearly cuts across the foliation of the amphibolite. 37 the early structures are associated with syntectonic aplites and pegmatites that cut across the amphibolite but occasionally occupy shear zones or fold hinges. the supracrustal series stratigraphy the supracrustal rocks that define the synform occur as largely discontinuous layers (figs 2–4), that could either correspond to an original, discontinuous geometry (therefore suggesting lava flows), or simply be a result of tectonic stretching during the multiphase deformation witnessed by the area. indeed, some of the contacts between the lithological units appear to be tectonic (figs 4, 5), suggesting that the present-day ‘stratigraphy’ might not be original. nevertheless, our mapping suggests that three main units can be recognised, allowing the following tentative stratigraphic sequence (figs 2, 4). 1. the lowermost, c. 100 m thick part consists of an association of amphibolite interlayered with garnet-sillimanite metapelites, sometimes with augen textures. some of the amphibolites are garnet-bearing, while others contain centimetre-sized lenses of diopside-bearing gabbro and small ultrabasic layers (pyroxenite or serpentinite, observed in the south-western part of the synform). the pelitic rocks seem to be more abundant in the northern limb and north-eastern extremity of the synform, while the ultramafic rocks and gabbros were found only in its south-western part. 2. the middle part is a sequence about 100 m thick of layered biotite gneiss, i.e. quartzo-feldspathic gneiss with no discriminant minerals and a compositional layering at a scale of c. 10 cm. the layered biotite gneiss is commonly interstratified with layers and lenses of amphibolite 10–100 cm thick. the contact with the lower amphibolite is gradational. as will be discussed below, the layered biotite gneiss likely represents meta-rhyolite. the middle unit of layered biotite gneiss probably does not have a constant thickness; furthermore, in poor, inland outcrops, it is readily confused with basement orthogneisses. a detailed log of the lower and middle parts of the sequence as described in the foregoing was made in the overturned, north-eastern part of the synform, displaying its complex and composite nature (fig. 3a, locs 81–82). 3. a horizon 50–100 m thick of fine grained, dark, layered amphibolite forms the highest observed level. the southern limb (locs 86–92) northern limb (locs 81–83) layered amphibolite intrusive pink granite possible peripheric intrusion? tectonic contact ? possible leucocratic intrusion? layered biotite gneiss (metarhyolite ?) occasional amphibolite layers layered amphibolitelayered amphibolite (garnet-bearing in places)(garnet-bearing in places) layered amphibolite (garnet-bearing in places) orthogneiss basement garnet-sillimanite metapelite interlayered with amphibolite. augen texture locally unconformity or tectonic contact? unconformity or tectonic contact? unconformity or tectonic contact? 10 cm lenses of diopsidebearing gabbro ultramafic layer amphibolite and layered grey gneisses interstratified b. generalised stratigraphic columns in the kangaatsiaq syncline a. coastal section, locs 81–82 sw ne loc. 82 loc. 81 538 537 536 535 533 532 530 100 m fig. 4. stratigraphic succession of the kangaatsiaq synform. a: detailed section of the overturned northern limb of the synform in its eastern extremity (locs 81–82), with sample numbers (all with prefix ‘485’). b: inferred generalised logs in the north-eastern and south-western parts of the synform. legend: see fig. 2. 38 a b c photo jfm-2002-4-14 photo jfm-2002-4-13 photo jfm-2002-4-12 photos jfm-2002-4-06–08 photos jfm-2002-4-06–08 photos jfm-2002-4-06–08 kangaatsiaq granite amphibolite layered biotite gneiss of supracrustal origin basement orthogneiss fig. 5. photomosaic (a) and structural interpretation (b) of the cliff face at loc. 80 (photo facing east). stippled: pegmatites; grey: high-strain zones. evidence for preto syn-granite, apparently extensional deformation is preserved in the amphibolite bodies intruded by the granite. details of the cliff face display the apparently extensive deformation in the amphibolite. cross-cutting pegmatites (see photo 4-12) are occasionally affected by this deformation, suggesting that it is synchronous or nearly synchronous with granite emplacement. c: schematic relationships between the granite, the early extensional deformation, and the supracrustal pile, inspired from loc. 80. 39 a1 a2 b1 b2 c1 c3 c2 c4 c5 fig. 6. field and thin sections photographs of lithologies of the supracrustal series (xpl: crossed polarised light; ppl: plane polarised light). microphotographs are c. 5 mm across. a1: outcrop of sillimanitebearing metapelite, loc. 64 (sample 485525). hammer is 80 cm long. a2: thin section (xpl) of the same. b1: outcrop of layered biotite gneiss interstratified with amphibolite at loc. 81 (sample 485537). pen is 15 cm long. b2: thin section (xpl) of same. c1: outcrop of the top amphibolite at loc. 58 (sample 485523). pocket knife is 10 cm long. c2: thin section (ppl) of same. c3: outcrop of gabbroic inclusions in the basal amphibolite layer at loc. 92 (sample 485541). compass 5 cm wide. c4: thin section (xpl) of clinopyroxene cluster in amphibolite. c5: thin section (xpl) of sample 485540 (ultramafic layer, same locality). 40 upper boundary of this unit is not observed, since it is everywhere intruded by the granite. this ‘top amphibolite’ is continuous and can be traced all around the exposed granite contact; it is also rather homogeneous, much more so than any of the other components of the supracrustal sequence. in loc. 80 (fig. 5), it appears to be in tectonic contact with the underlying layered biotite gneiss. field description and petrology as mentioned above, three main components are observed in the supracrustal succession: aluminous metapelite, layered biotite gneiss and amphibolite. field aspects together with photographs of thin sections are presented in fig. 6. the aluminous metapelites occur as slaty, fine-grained (0.5–1 mm), grey to yellowish paragneisses (fig. 6a1). garnet or sillimanite is commonly seen in outcrop. in thin section, they display biotite, plagioclase and quartz with either sillimanite or poikiloblastic garnet (fig. 6a2) cutting across an earlier weak foliation marked by preferred orientation of biotite flakes and elongation of plagioclase crystals. the layered biotite gneisses appear as grey, relatively massive, fine grained (0.5–1 mm), finely layered rocks. they are interstratified at all scales with amphibolite (figs 3c, 6b1) and generally form discontinuous bodies on a 100 m scale. they consist of quartz, plagioclase, k-feldspar and biotite; the foliation is defined by the preferred orientation of biotite and elongation of quartz grains (fig. 6b2). the amphibolites are dark, massive rocks that also show a strong compositional banding (fig. 6c1–c2). regardless of their mode of outcrop either as a thick continuous layer, as in the ‘top amphibolite’, or as discontinuous layers interstratified with other lithologies, they are very similar in visual aspect and mineralogy. they mostly consist of a fine-grained (0.5–1 mm) hornblende-plagioclase assemblage, with preferred orientation of minerals defining the foliation. commonly, small clusters of clinopyroxene surrounded by felsic (mostly plagioclase) rims are observed (fig. 6c4). at one locality, gabbroic lenses on a scale of 5–10 cm have been observed within the amphibolite (loc. 92, fig. 6c3). they are medium grained (2–5 mm) and greenish in aspect, and composed of a clinopyroxene-plagioclase association with diffuse contacts with the neighbouring amphibolite (fig. 6c4). at the same locality, an ultramasiosio2 2 al ka lin e subalkaline/tholeiitic subalkaline/tholeiitic subalkaline/tholeiiticsubalkaline/tholeiiticsubalkaline/tholeiitic 40 50 60 70 80 0 5 10 15 ● ● paragneisses paragneisses paragneiss na2o + k2o basement supracrustal sequence kangaatsiaq granite layered biotite gneiss layered amphibolite ultramafic rocks ●‘normal’ orthogneiss high-k orthogneiss amphibolite enclaves basaltic andesite dacite rhyolite basalt basalt basalt andesite siosio2 2 sio2 fig. 7. total alkali vs. silica (tas) diagram (le maître et al. 1989) for the magmatic components of the supracrustal rocks and the surrounding orthogneisses. 41 fic layer c. 0.5 m thick has been observed. it is slightly coarser grained (2–5 mm) than the amphibolite, and solely consists of amphibole grains (fig. 6c5), which are optically similar to the hornblende in the surrounding amphibolite. geochemistry and origin figures 7–8 and table 1 summarise the major and trace element (especially ree) characteristics and relationships of the three main supracrustal components: amphibolites, metapelites and layered biotite gneisses. there is little, if any doubt of the fact that the amphibolites correspond to metamorphosed and deformed mafic igneous rocks. elsewhere, similar field characteristics in amphibolites as those observed here have been interpreted as corresponding to transposition of former pillow lavas in high strain domains (e.g. myers 2001). the metapelites obviously have a sedimentary origin and probably represent terrigeneous sediments. the origin of the layered biotite gneisses, however, is less obvious. they could represent either sedimentary or felsic volcanic rocks. therefore, they are plotted on geochemical diagrams for both magmatic and sedimentary rocks (see below), allowing comparisons. origin of the amphibolites the supracrustal amphibolites and their counterparts, enclaves in the basement orthogneisses, appear to be very similar in composition. they plot mostly as basalts in a tas diagram (fig. 7; le maître et al. 1989), and an afm diagram (fig. 9; irvine & baragar 1971) reveals that they belong to a tholeiitic series. this, together with their spectacularly flat ree pattern at about 10 times chondritic values (fig. 8), is consistent with the amphibolites corresponding to former morb basalts, possibly formed as part of an oceanic crust. many discriminant diagrams for basaltic rocks have been proposed on geochemical grounds (e.g. pearce 1982; shervais 1982; mullen 1983). however, some caution should be exercised when using such diagrams for the archaean, since the existence of modern-style tectonic settings in the archaean is not certain, and the palaeogeodynamical contexts might not be similar to those of modern settings (hamilton 1998; mccall 2003; van kranendonk 2003). nevertheless, in 100 10 1 0.1 basement orthogneiss la ce pr nd pm sm eu gd tb dy ho er tm yb lu supracrustal amphibolite amphibolite as enclaves in orthogneisses amphibolite100 10 1 0.1 la ce pr sa m pl e / r ee c ho nd ri te layered biotite gneiss100 10 1 0.1 la ce pr nd pm sm eu gd tb dy ho er tm yb lu sa m pl e / r ee c ho nd ri te kangaatsiaq granite la ce pr nd pm sm eu gd tb dy ho er tm yb lu nd pm sm eu gd tb dy ho er tm yb lu 1000 100 10 1 0.1 aluminous metapelites100 10 1 0.1 sa m pl e / r ee c ho nd ri te sa m pl e / r ee c ho nd ri te sa m pl e / r ee c ho nd ri te la ce pr nd pm sm eu gd tb dy ho er tm yb lu fig. 8. ree patterns (chondrite normalised, boynton 1984) for the lithologies in and around the kangaatsiaq synform. 42 such diagrams, the amphibolites plot either as morb or as rocks originated in oceanic arcs (arc tholeiites), leaving some ambiguity about their original setting. origin of the aluminous metapelites the geochemistry of metasedimentary rocks is commonly used to discuss their source, in terms of (1) the nature of the original sediments, (2) the nature of the weathered/ eroded source material, and (3) the degree of weathering of the source (see e.g. taylor & mclennan 1985; herron 1988; roser & korsch 1988; nesbitt & young 1989; bohlar et al. 2005). while several authors also use the geochemistry of sediments to discuss their geodynamical setting (bhatia 1983; bhatia & crook 1986; roser & korsch 1988), some caution should be exercised when dealing with archaean environments, as mentioned above. in terms of classification, the metasedimentary rocks from the kangaatsiaq area plot mostly as shales or greywackes, using either of the two schemes proposed by herron (1988). one of these is shown on fig. 10a; the ambiguity and possible (chemical) confusion between the two groups, shales and greywackes, which are poorly separated by this diagram, has been outlined by these authors. nevertheless, the conclusion points to relatively immature sediments which have undergone limited transport from their source. the nature of the source itself can be discussed using major or trace elements. roser & korsch (1988) proposed a scheme for source determination of clastic sediments on the basis of major elements. in this instance, the studied samples straddle the p2–p3 boundary (fig. 10b), suggesting a felsic to intermediate source. also trace elements can be used to refine this conclusion. as pointed out by taylor & mclennan (1985), some elements (high field strength elements, rare earth elements (ree), y, sc, th) only undergo limited fractionation during sedimentary processes; thus, their ratios reflect the signature of their source. plotting these elements against each other shows that the kangaatsiaq metasedimentary rocks (fig. 10e– h) have element ratios that are generally consistent with derivation from an orthogneissic source (amphibolites generally have too low trace element contents and incorrect ratios to be a possible source). the only exception is for heavy ree (figs 8, 10g). indeed, the relatively high yb contents of the metasedimentary rocks precludes their derivation solely from a low-yb gneissic basement, and implies that they must, at least in part, have been derived from higher-yb rocks such as the amphibolites; this is hardly a surprise, since amphibolite occurs as enclaves intercalated within the orthogneisses. modelling the ree contents of such a mixture shows that mixing of orthoa m f ● ● granite orthogneiss tholeiite series calc-alkaline series supracrustal amphibolite amphibolite enclave ultramafic rocks fig. 9. afm diagram (irvine & baragar 1971) showing the tholeiitic affinity of both the supracrustal amphibolites and the enclaves in the gneisses. a, na 2 o + k 2 o; f, feo total ; m, mgo. the fields of the basement orthogneisses and the kangaatsiaq granite are also shown for comparison. facing page: fig. 10. major and trace element geochemistry (a–d and e–h) of the metasedimentary rocks (paragneisses, and layered biotite gneisses). dotted fields show the compositions of the major regional lithologies (orthogneiss and high-k orthogneiss, moyen et al. 2003a; steenfelt et al. 2005; amphibolite; kangaatsiaq granite). a: log(sio 2 /al 2 o 3 ) vs. log(fe 2 o 3 /k 2 o), from herron (1988). b: discriminant diagrams for the metapelites, from roser and korsch (1988). the sources for each group are p1, mafic to intermediate volcanic rocks; p2, intermediate (andesitic, dacitic, occasionally rhyolitic) volcanic rocks; p3, felsic volcanic rocks; p4, evolved sediments, sandstones, etc. the discriminant functions are: f1 = –1.773 tio 2 + 0.607 al 2 o 3 + 0.760 fe 2 o 3 – 1.500 mgo + 0.616 cao + 0.509 na 2 o – 1.224 k 2 o – 9.090; f2 = 0.445 tio 2 + 0.070 al 2 o 3 – 0.250 fe 2 o 3 –1.142 mgo + 0.438 cao + 1.475 na 2 o + 1.426 k 2 o – 6.861. c, d: triangular diagrams (from nesbitt & young 1989). stars: theoretical mineral compositions; il, illite; ms, muscovite; pg, plagioclase; ksp, k-feldspar; cpx, clinopyroxene; hbl, hornblende; chl, chlorite; bt, biotite; sm, smectite. dashed arrows: trends for (1) weathering and (2) k-metasomatism, after nesbitt & young (1989) and bohlar et al. (2005). e, f: u vs. th and ti vs. zr (log scale) diagrams, showing that the metasedimentary rocks have trace elements ratios comparable to the gneisses, but mostly different from the amphibolites. g, h: la/yb vs. yb and ti/zr vs. ni (log scale) diagrams displaying the same relationships as e–f, also showing the mixing between an amphibolite-like and an orthogneiss-like source (ticks at 10% increments). 43 amphibolite yb la/yb ti/zr u ti ni zrth 0 5 10 0.0 0.5 1.0 1.5 th/u = 1 th/u = 5 th/u = 10 orthogneiss amphibolite amphibolite amphiboliteamphiboliteamphibolite granite 10 100 200 100 200 500 10 20 50 1000 2000 5000 10000 0 20 40 60 80 100 120 140 20 50 100 200 ti/zr = 10 ti/zr = 50 ti/zr = 100 ti/zr = 200 ti/zr = 20 orthogneissorthogneissorthogneissorthogneissorthogneiss amphibolite granite 0.0 0.5 1.0 1.5 2.0 2.5 3.0 orthogneisses amphibolite granite 0.1 0.5 1 5 10 50 100 500 orthogneiss amphibolite granite e f f f g h amphibolite aluminous paragneiss layered biotite gneiss standard mineral compositions for reference compositional fields of regional lithologies (see figure text) granite chl ksp, pg sm il bt cpx hbl ms pg ksp il, ms hbl cpx al2o3 k2o al2o3 cao + na2o + k2ocao + na2o feot + mgo 1 2 1 2 1 orthogneissorthogneiss c d granitegranitegranite 0.0 0.5 1.0 1.5 2.0 -0.5 0.0 0.6 1.0 1.5 fe -sh ale sh ale fe -sa nd su bli tha ren ite qua rtz are nit e su ba rko se ar ko se w ac ke lit ha ren ite lo g (f e 2 o 3/ k 2o ) log(sio2/al2o3)log(siolog(sio2/al/al2o3)log(sio2/al2o3) amphiboliteamphiboliteamphibolite orthogneiss granite a p4 p3 p2 p1 8 4 0 -4 -8 8 4 0 -4 -8 f1 f2 amphiboliteamphibolite orthogneissorthogneiss amphibolite orthogneiss granite b 44 si o 2 t io 2 a l 2 o 3 fe 2o 3* m no m go c ao n a 2 o k 2o p 2 o 5 h 2o k /n a m g# a /c n k c .i. a . v c r n i c o c u z n r b sr ba y z r n b ta h f sc g a c s pb t h u t h/ u t i/z r la c e pr n d sm eu g d t b d y h o er t m y b lu la /y b eu /e u* 69 .0 0 74 .0 8 72 .9 2 72 .8 8 70 .2 1 66 .0 4 0. 37 0. 18 0. 27 0. 21 0. 43 0. 62 15 .1 6 13 .2 2 13 .4 8 13 .8 7 15 .1 4 16 .0 1 2. 46 1. 33 1. 48 1. 61 2. 50 3. 94 0. 02 0. 01 0. 01 0. 02 0. 04 0. 01 0. 79 0. 31 0. 44 0. 46 0. 89 1. 27 2. 07 1. 01 1. 17 1. 29 2. 17 5. 07 4. 18 3. 63 3. 64 4. 06 3. 93 3. 31 4. 23 4. 56 4. 74 4. 28 4. 29 1. 07 0. 16 0. 05 0. 07 0. 08 0. 17 0. 39 0. 22 0. 16 0. 17 0. 26 0. 80 0. 67 0. 83 0. 86 0. 69 0. 72 0. 21 39 32 37 36 41 39 1. 00 1. 04 1. 02 1. 02 1. 01 1. 01 28 11 10 11 33 36 7 3 3 4 14 11 8. 1 3. 5 4. 4 5. 5 bd l 22 .7 37 .0 44 .6 79 .4 68 .8 50 .7 3. 3 2. 6 2. 6 1. 7 13 .0 32 9. 1 36 .2 21 .5 20 .6 27 .9 25 .0 31 .9 11 4. 7 11 5. 5 11 8. 7 12 5. 8 56 .0 96 .0 73 5 25 2 34 2 34 0 71 2 19 13 14 76 46 3 67 2 59 3 13 84 20 95 8. 5 2. 7 5. 7 5. 5 bd l 7. 2 17 8 10 4 13 2 14 0 21 9 20 3 8. 4 3. 8 6. 0 4. 8 bd l 7. 0 4. 2 3. 1 3. 5 3. 7 4. 6 5. 4 2. 2 10 .5 2. 6 4. 7 2. 4 0. 9 1. 5 1. 5 3. 9 19 .1 17 .6 17 .8 18 .2 21 .3 1. 0 0. 4 0. 6 1. 9 6. 2 15 .6 16 .5 12 .8 17 .3 6. 9 10 .8 6. 8 9. 8 9. 2 14 .0 0. 9 1. 4 1. 0 0. 9 2. 6 12 .1 5. 0 10 .0 9. 8 5. 4 12 .6 10 .4 12 .3 9. 0 11 .8 18 .2 35 .9 19 .8 35 .2 39 .1 29 .0 15 1. 7 81 .2 39 .7 68 .9 75 .1 10 5. 0 31 3. 7 9. 7 4. 2 7. 5 8. 2 37 .7 33 .0 12 .8 23 .3 25 .4 12 6. 4 5. 3 1. 6 3. 2 3. 1 13 .4 1. 5 0. 6 0. 7 0. 7 2. 4 5. 1 1. 9 3. 0 3. 2 12 .5 0. 44 0. 15 0. 27 0. 27 0. 83 1. 96 0. 62 1. 14 1. 10 2. 81 0. 28 0. 09 0. 17 0. 17 0. 25 0. 76 0. 27 0. 50 0. 49 0. 98 0. 10 0. 03 0. 07 0. 06 0. 07 0. 57 0. 24 0. 45 0. 43 0. 47 0. 09 0. 04 0. 06 0. 06 0. 06 63 .0 83 .3 79 .0 89 .8 32 0 0. 87 0. 96 0. 70 0. 71 0. 56 62 .2 9 70 .9 7 63 .1 0 66 .0 3 0. 68 0. 26 0. 61 0. 62 15 .6 9 14 .8 8 16 .9 7 18 .9 8 5. 68 1. 85 5. 92 4. 27 0. 09 0. 02 0. 10 0. 03 3. 13 0. 48 2. 17 1. 45 4. 71 2. 13 4. 60 2. 19 3. 31 4. 23 2. 94 3. 20 2. 03 3. 38 1. 41 1. 70 0. 22 0. 08 0. 07 0. 10 0. 98 0. 18 0. 69 0. 59 0. 40 0. 53 0. 32 0. 35 52 34 42 40 0. 97 1. 03 1. 15 1. 71 60 .9 5 60 .4 4 65 .4 7 72 .8 0 1 00 12 13 3 88 68 2 13 6 61 50 .1 5. 0 79 .1 25 .1 47 .8 86 .3 10 0. 8 10 3. 0 14 .2 3. 7 10 1. 4 14 .9 67 .0 34 .7 78 .3 57 .3 46 .9 83 .2 44 .7 79 .8 8 03 82 2 23 1 21 3 8 19 14 30 31 4 26 2 16 .2 3. 8 12 .5 9. 2 1 65 14 7 98 93 8. 3 3. 5 4. 5 5. 1 3. 9 3. 8 2. 7 2. 4 3. 6 3. 5 3. 3 2. 3 14 .2 2. 3 23 .4 13 .0 18 .6 19 .6 19 .7 24 .1 4. 7 1. 2 2. 2 1. 5 8. 6 15 .4 7. 1 7. 8 4. 7 4. 8 2. 7 3. 2 0. 8 0. 6 0. 8 0. 7 5. 6 8. 2 3. 3 4. 7 24 .8 10 .5 37 .0 39 .7 35 .3 27 .3 12 .9 21 .3 74 .4 61 .0 27 .2 46 .0 9. 3 6. 9 3. 4 5. 8 33 .4 22 .2 13 .0 21 .5 4. 9 2. 8 2. 5 3. 4 1. 3 0. 9 0. 7 0. 9 5. 0 2. 8 2. 6 3. 1 0. 57 0. 22 0. 37 0. 36 2. 99 0. 86 2. 13 1. 86 0. 53 0. 13 0. 43 0. 33 1. 51 0. 35 1. 20 0. 93 0. 22 0. 05 0. 18 0. 14 1. 41 0. 30 1. 23 0. 95 0. 22 0. 04 0. 18 0. 14 25 .0 91 .8 10 .5 22 .3 0. 77 0. 93 0. 87 0. 82 70 .8 1 70 .0 3 47 .5 2 0. 29 0. 31 0. 74 15 .0 8 15 .1 5 14 .9 4 2. 47 2. 06 12 .4 6 0. 02 0. 02 0. 19 0. 90 0. 64 8. 32 2. 88 2. 39 11 .2 2 4. 42 4. 61 2. 48 2. 02 3. 13 0. 83 0. 06 0. 15 0. 04 0. 18 0. 26 0. 54 0. 30 0. 45 0. 22 42 38 57 1. 03 0. 99 0. 59 25 20 25 4 22 3 33 2 12 .9 7. 6 14 7. 8 71 .4 46 .4 72 .3 5. 4 3. 2 40 .3 43 .6 42 .7 86 .2 74 .4 92 .0 7. 7 50 4 83 9 93 55 8 10 49 39 3. 0 4. 5 17 .6 86 2 11 7 33 4 5. 4 5. 6 2. 1 2. 5 3. 2 1. 0 11 .4 7. 9 3. 1 5. 0 3. 7 43 .8 20 .8 22 .4 15 .1 1. 0 1. 3 0. 0 9. 6 15 .6 2. 6 1. 9 5. 1 0. 2 0. 5 1. 7 0. 3 3. 8 3. 0 0. 7 20 .4 15 .9 13 4 10 .2 38 .4 2. 6 21 .4 82 .7 6. 3 2. 3 8. 9 1. 0 7. 8 30 .1 4. 9 1. 3 3. 8 1. 6 0. 6 1. 0 0. 6 1. 4 3. 9 2. 3 0. 13 0. 30 0. 43 0. 69 1. 35 2. 74 0. 11 0. 16 0. 63 0. 32 0. 52 1. 68 0. 04 0. 06 0. 28 0. 27 0. 38 1. 82 0. 03 0. 05 0. 26 37 .8 10 2 1. 44 1. 34 0. 78 0. 94 47 .0 5 47 .1 4 48 .7 7 48 .7 4 48 .0 8 45 .8 7 0. 95 0. 81 0. 87 0. 77 0. 72 0. 51 14 .6 8 14 .7 2 14 .6 7 15 .8 2 14 .6 9 11 .5 4 12 .9 9 10 .9 0 12 .6 0 12 .1 4 10 .7 7 11 .7 4 0. 22 0. 26 0. 24 0. 21 0. 20 0. 21 6. 80 4. 74 6. 70 7. 18 7. 90 14 .1 9 13 .3 0 17 .8 9 11 .5 1 12 .1 8 13 .1 4 11 .7 8 2. 02 1. 40 2. 67 1. 64 2. 55 1. 39 0. 53 0. 08 0. 76 0. 31 0. 52 0. 58 0. 06 0. 06 0. 06 0. 05 0. 03 0. 04 0. 36 1. 19 0. 51 0. 45 0. 30 0. 79 0. 17 0. 04 0. 19 0. 12 0. 13 0. 27 51 46 51 54 59 71 0. 52 0. 42 0. 56 0. 63 0. 51 0. 47 27 1 23 7 27 1 24 4 22 5 17 7 20 7 30 8 28 4 24 7 42 5 17 34 15 3. 1 16 7. 6 12 1. 7 17 5. 7 19 7. 5 72 8. 1 72 .1 59 .9 57 .1 57 .5 61 .0 81 .6 78 .7 73 .7 30 .3 12 .6 24 .1 2. 2 90 .1 76 .7 10 8. 2 77 .4 74 .8 10 3. 3 9. 1 1. 7 9. 4 32 .4 6. 5 14 .8 93 11 7 11 7 92 94 78 64 22 11 8 46 37 82 21 .1 18 .8 20 .6 18 .5 17 .1 11 .6 19 16 20 32 16 24 2. 8 2. 2 2. 4 4. 1 2. 7 1. 8 0. 9 0. 6 0. 9 1. 1 0. 7 0. 8 1. 1 0. 8 0. 8 0. 7 1. 3 0. 6 41 .6 38 .2 40 .9 41 .2 38 .3 31 .3 16 .4 14 .5 16 .1 16 .5 14 .8 11 .9 0. 1 0. 1 0. 0 0. 3 0. 0 0. 3 2. 9 2. 8 3. 5 1. 7 3. 2 1. 3 0. 4 0. 2 0. 3 0. 2 0. 2 0. 2 0. 2 0. 1 0. 1 0. 3 0. 1 0. 3 1. 8 3. 5 3. 2 0. 9 1. 4 0. 5 29 6 30 3 26 6 14 2 27 0 12 5 3. 0 2. 2 3. 0 2. 2 2. 9 2. 6 7. 7 6. 2 7. 2 6. 1 7. 9 5. 2 1. 2 1. 0 1. 2 1. 0 1. 3 0. 8 6. 3 5. 4 6. 1 4. 9 6. 1 3. 7 2. 1 1. 8 2. 0 1. 8 1. 8 1. 2 0. 7 0. 6 0. 7 0. 7 0. 6 0. 4 2. 5 2. 1 2. 5 2. 3 2. 1 1. 4 0. 49 0. 43 0. 48 0. 45 0. 41 0. 28 3. 12 2. 85 3. 10 2. 89 2. 60 1. 72 0. 72 0. 64 0. 73 0. 64 0. 60 0. 40 1. 96 1. 77 1. 95 1. 72 1. 62 1. 13 0. 31 0. 28 0. 33 0. 27 0. 26 0. 18 2. 12 1. 85 2. 08 1. 82 1. 70 1. 20 0. 31 0. 29 0. 31 0. 26 0. 26 0. 17 1. 42 1. 22 1. 43 1. 20 1. 68 2. 19 0. 96 1. 01 0. 90 1. 00 0. 87 0. 99 la ye re d bi ot ite g ne is s pa le le uc o cr at ic g ne is s bi ot ite -b ea ri ng g ne is s a m ph ib ol ite k an ga at si aq g ra ni te a lu m in ou s m et as ed . 48 55 31 48 55 35 48 55 25 q zbt -s ill -g t sc hi st q zbt -g t gn ei ss 81 81 81 81 52 64 48 55 37 48 55 38 48 55 23 48 55 33 48 55 36 48 55 28 48 55 40 81 58 73 92 61 73 75 85 gi a 20 01 -2 17 81 a m ph ib ol ite in lo w er la ye r as so ci at ed w ith fe ls ic r oc ks pe ri ph er ic gr an iti c dy ke m as si ve a m ph i bo lit e, t op u ltr am af ic la ye r in a m ph ib ol ite 48 55 29 48 55 39 48 55 24 48 55 27 47 05 29 48 55 34 81 81 57 48 55 22 48 55 30 48 55 32 ba se m en t m ai n gr an iti c m as s o rt ho gn ei ss a m ph ib ol ite en cl av e sa m pl e n o. lo ca lit y ta bl e 1. c he m ic al a na ly se s of r oc ks in a nd a ro un d th e k an ga at si aq s yn cl in e sa m pl e nu m be rs r ef er t o g eu s da ta ba se s; lo ca lit ie s ar e sh ow n on f ig s 1 an d 4, e xc ep t gi a 20 01 -2 17 o ut si de t he m ap a re a (u t m 3 96 54 0, 7 57 59 84 ). m aj or e le m en ts in w t% ; t ra ce e le m en ts in p pm . a na ly tic al d at a ob ta in ed a t g eu s by x r f (m aj or e le m en ts a nd a pp ro xi m at e tr ac e el em en ts in 4 70 52 9) a nd ic pm s (a ll ot he r tr ac e el em en ts ). k /n a: m ol ec ul ar k /n a ra tio . m g# : m ol ec ul ar 1 00 m g/ (m g + f e) . a /c n k : m ol ec ul ar a l/( c a + n a + k ). c .i. a .: c on tin en ta l i nd ex o f a lte ra tio n (n es bi tt & y ou ng 1 98 9) ; m et as ed .: m et as ed im en ta ry r oc k; b dl : b el ow d et ec tio n lim it. 45 gneisses with amphibolite (fig. 10g, h) can explain the yb contents of the sediments; since the amphibolites are, collectively, less enriched in trace elements than the gneisses, their involvement would only have little effect on the other incompatible elements in the metasedimentary rocks. in contrast, the latter display higher ni and cr contents than the orthogneisses, also consistent with a contribution from amphibolite or its precursor rocks in their formation (fig. 10h). finally, the degree of alteration of the source can be discussed. the metasedimentary rocks display c.i.a. values (chemical index of alteration, nesbitt & young 1989) of c. 60–70 (table 1), slightly lower than for shales or similar rocks (70–75, taylor & mclennan 1985). in the triangular diagrams proposed by nesbitt & young (1989; fig. 10c, d), they also depart only moderately from their protoliths, suggesting a relatively unweathered source. very little or no evidence for secondary kenrichment is observed. taking the above-mentioned limitations into account, the geoynamic setting inferred from the geochemistry gives consistent results regardless of the classification scheme used. both the major elements classifications of bhatia (1983) and roser & korsch (1988) and the trace element systems of bhatia & crook (1986) suggest an oceanic or continental island-arc setting. however, this only reflects the characteristics outlined above: relatively immature sediment derived from poorly weathered felsic to intermediate magmatic rocks, with a possible mafic component. origin of the layered biotite gneiss (felsic volcanic rocks?) the two samples analysed of the layered biotite gneisses give ambiguous geochemical signatures and can be interpreted either as sedimentary or igneous (figs 7–10). in general, they seem to share more similarities with the granite or the orthogneisses than with any other member of the supracrustal group. in particular, fig. 10 (c, d) shows that if these rocks are of sedimentary origin, they are indeed very similar to their source and were derived from a largely unweathered protolith. this implies that the layered biotite gneiss can be interpreted in two ways. it may represent very immature sediment derived from a mostly unweathered protolith with a very similar bulk composition, such as a conglomerate made of pebbles of unweathered orthogneiss, in which case the banding could be a trace of the transposed pebbles. alternatively the layered biotite gneiss represents calc-alkali or ttg-type felsic lavas, whose composition would of course be very similar to that of their plutonic counterparts. origin of the supracrustal sequence as a whole based on the foregoing discussion two interpretations can be proposed for the supracrustal sequence. 1. the succession could represent a dismembered ophiolite sequence intermingled with clastic sediments eroded from a nearby continent. the combined sequence could then be interpreted as an accretionary prism. the likely tectonic nature of the contact between members of the sequence (see above and figs 4, 5) supports this hypothesis. 2. the whole supracrustal pile consists of a bimodal, calcalkaline, probably subduction-related volcanic suite associated with immature terrigeneous sediments directly derived from their weathering. this is consistent with an arc situation, in which a backor fore-arc basin is being filled with both volcanic products and detrital sediments largely derived from the weathering of these lavas. at the same time, tonalitic plutons are emplaced at depth from the same magmas. the plutonic rocks are quickly uplifted and eroded, and, besides intruding into the supracrustal pile, may in some cases also represent the basement for subsequent volcano-detritic basin fill. in both cases, the rocks were formed in a convergent setting, probably above or close to an active subductions margin. in general, arcor subduction-related origins for archaean volcanic suites are preferred by most workers (e.g. card 1990; lowe 1994; windley 1995; chadwick et al. 1996), although the issue remains controversial (hamilton 1998; mccall 2003; van kranendonk 2003). nature and origin of the kangaatsiaq granite the kangaatsiaq granite is pink, porphyritic, and displays a distinct rodding (fig. 11a) showing that it has been intensely deformed (see below). while yz sections (perpendicular to the main stretching direction) display a preserved magmatic texture, sections parallel to x clearly show the gneissic texture of the rock. the mineralogical paragenesis is k-feldspar + quartz + sodic plagioclase + biotite, with accessory zircon, titanite, apatite and oxides. the granite has moderate k/na ratios (0.67–0.86), is slightly metaluminous with a/cnk ratios of 1.00–1.04, and has low mg# of 32– 41 (table 1). ni and cr contents are also low, while rb, sr and ba contents are moderate; 46 this composition corresponds to the biotite-bearing granites of moyen et al. (2003b), which are interpreted to have been derived from partial melting of ttg gneisses. this conclusion is consistent with the highly migmatitic nature of the surrounding gneissic basement (van gool et al. 2002a). structure and deformation history as mentioned above, the granite displays a strong rodding and l > s fabrics (fig. 11a). the strain pattern in the granite (fig. 11b) is consistent with highly constrictional deformation, with foliation poles plotting on a great circle, and lineations clustered near the pole of this great circle. this corresponds to subhorizontal, ene–wsw stretching, consistent with the general orientation of the structures in kangaatsiaq area (fig. 1), and more generally with the structural grain of the region (van gool et al. 2002a; piazolo et al. 2004; mazur et al. 2006, this volume). the surrounding gneissic basement and supracrustal rocks show the same strain pattern when plotted (fig. 11c), although in the field, the rocks commonly have a ls or s > l fabric. this suggests that pre-existing foliations have been reoriented during the latest constrictional deformation event, leading to their present distribution. the fact that intense constriction (rather than shortening) can produce folded structures has previously been demonstrated by e.g. leloup et al. (1995) in the red river shear zone in yunnan, china, where the ductile deformation in gneisses resulted in the development of elongate synclines and anticlines with axes parallel to the shear zone and the x-axis of deformation. the study of deformation-related textures allows the conditions of deformation to be roughly constrained. in the granite and felsic components of the supracrustal sea total data: 70 equal area, lower hemisphere b. within the granite poles to foliation lineation poles to axial planes fold axis c. outside the granite total data: 42 fig. 11. a: macroscopic view of the kangaatsiaq granite at loc. 75 (corresponding to sample 485529), showing strong rodding. hammer shaft about 4 cm wide. b, c: stereograms of poles to foliation (circles) and lineations (squares) within and outside the granitic intrusion. the strain patterns are similar in both units and define a highly constrictional, ne–sw-trending and subhorizontal deformation. 47 a b c d e fig. 12. deformation textures either related to the latest, constrictional deformation (a–c) or not compatible with low-t deformation (d, e). see comments in the main text. a: quartz ribbons in the kangaatsiaq granite (sample 485527). b: quartz subgrains in felsic supracrustal gneiss (sample 485531). c: poikiloblastic garnet in metapelite cutting across an earlier foliation (sample 485535). d, e: high-temperature recrystallisation with 120° triple junctions in amphibolite (sample 485540) and felsic rocks (sample 485530). in e, the quartz also shows low-temperature deformational features such as undulating extinction and quartz subgrains, indicating that this rock witnessed two successive deformation events. 48 ries (fig. 12a, b), the deformation led to the development of quartz subgrains and recrystallised quartz ribbons. this corresponds to deformation under lower amphibolite facies conditions (400 ± 50°c; bouchez & pécher 1976; gapais & barbarin 1986; gapais 1989; hirth & tullis 1992; vernon 2000). under these conditions, only the quartz is ductile, such that all deformation is accommodated by quartz recrystallisation or deformation. in the al-rich lithologies, deformation-related textures are mostly seen in the formation of poikiloblastic, synto post-tectonic garnets (fig. 12c). piazolo (2002) estimated that the chemistry of garnet in similar pelites nearby is compatible with a long duration of temperature conditions at around 500°c, which is in broad agreement with the above estimate. willigers et al. (2002) described the cooling history of the nno close to our study area from ar-ar dating of various minerals, and likewise concluded that the cooling history of the nno was slow, from 400°c (muscovite closure) at 1.7 ga to 200°c (k-feldspar closure) at 1.5 ga. therefore, it can be considered that a constrictional deformation event post-dating the granite emplacement occurred during cooling to lower amphibolite facies conditions. since this event is apparently responsible for the regional-scale structures (mazur 2002; van gool et al. 2002b; piazolo et al. 2004), and is of lower proterozoic age (willigers et al. 2002), we propose that it corresponds essentially to the nagssugtoqidian deformation proper. however, some textures are not compatible with the above conditions. in amphibolites, high-temperature fabrics with polygonal textures and 120° triple junctions are preserved (fig. 12d). in some of the felsic supracrustal rocks or basement gneisses (but never in the granite), evidence is preserved for a similar high-temperature fabric, overprinted by later quartz recrystallisation (fig. 12e). according to kretz (1969), gower & simpson (1992), kretz (1994) and martelat et al. (1999), such fabrics are poikiloblastic garnet cutting the d1 fabric quartz subgrains, etc. quartz subgrains, etc. high-temperature recrystallisation country rocks granite pr en ag ss ug to qi di an d 1/ d 1b (g ra nu lit ic ) n ag ss ug to qi di an d 2 (a m ph ib ol iti c) (not formed) no quartz d2 deformation not expressed quartzo-feldspathic amphibolitic granite emplacement fig. 13. summary of the deformation history of the kangaatsiaq synform and kangaatsiaq granite. see comments in the main text. photos from fig. 12. 49 likely to develop under granulite facies conditions (600– 800°c). this points to the existence of one or more older (d1?) deformation event(s). since no evidence for this deformation is found in the granite, we suggest that it was pre-granite, and therefore likely corresponds to late archaean deformation. p–t estimates for metapelites and metabasites in the kangaatsiaq area by piazolo (2002) also indicated the existence of an early metamorphic phase with p–t conditions between 650°c, 3–5 kbar and 780°c, p unknown. this estimate is in good agreement with the textural evidence for d1 deformation under lower granulite facies conditions. the pre-granite deformation is also evidenced by the early isoclinal folds, the existence of a foliation within the supracrustal rocks that does not exist in the granite, and the fact that the granite apparently cuts earlier structures (fig. 5). at loc. 80, the granite is clearly observed cutting across the foliation and shear bands in the amphibolite; these shear bands are injected by pegmatites that might also be cut by the granite. this suggests that there were actually two pre-granite events, the first of which corresponds to the granulite facies formation of the foliation and isoclinal folds, and the second one to the pegmatite-injected shear bands. however, the floor of the granitic intrusion is also apparently offset by the shear bands (fig. 5b). furthermore, the geometry of the shear bands and the foliation suggests extensional deformation; since the cliff face studied here almost corresponds to a yz section relative to the regional constrictional deformation, this geometry is likely to correspond to the original, preserved pre-constriction geometry. finally, the fact that the granite both cuts across, and is offset by the shear bands, suggests that the granite emplacement may actually have been syn-extension as sketched in fig. 5c. altogether, the simplest possible deformation history (with the smallest number of episodes) can be summarised as follows (fig. 13). 1. a first deformation event (d1) under lower granulite facies conditions (c. 5 kb, 600–800°c), resulted in the development of granulitic (polygonal) textures in all the existing lithologies, the formation of a main foliation, and isoclinal folding. it probably corresponds to compression of the original, likely accretionaly wedge or arc sequence. 2. this may have been followed by a second event (d1b) of probably extensional deformation, maybe associated with (or shortly followed by) the emplacement of the granite sheet. this event, only witnessed by the shear zones cutting the d1 foliation, e.g. at locality 80, is poorly recorded and probably just represents the final stage of d1 deformation. assuming the granite has a late archaean age, which is very likely in the regional context, this deformation could correspond to the later stages of the evolution of an arc or active continental margin, with strain relaxation and syn-extension granite emplacement. 3. a final event of constrictional deformation under lower amphibolite conditions (d2). due to the relatively low-temperature conditions, only the quartz-bearing lithologies were affected. therefore, the granite shows strong recrystallisation, the felsic supracrustal rocks display overprinting of the d1/d1b fabric by this event, and the quartz-free amphibolites were essentially unaffected by this event. the d2 event corresponds to the purely constrictional, regional structures which have been interpreted by piazolo et al. (2004) and mazur et al. (2006, this volume) as resulting from the indentation of the nno by a solid, north-moving block immediately north of the arfersiorfik shear zone (for the latter, see e.g. sørensen et al. 2006, this volume). this palaeoproterozoic deformation gave the studied area its present synformal structure. conclusions while the present-day synclinal structure of the kangaatsiaq area essentially results from n60 constriction related to the palaeoproterozoic nagssuqtoqidian deformation, the lithologies together with early preserved structures give insight into the late archaean crustal evolution. the basement gneisses genetically belong to the ttg suite (moyen et al. 2003a; steenfelt et al. 2005), which is generally interpreted as generated by partial melting of a subducting slab (e.g. martin 1994). some components of the basement display implications of mantle wedge involvement in their genesis (steenfelt et al. 2005), which is unusual in the archaean but nevertheless consistent with an active margin setting. the supracrustal succession is composed of discontinuous layers of mafic morb-like or arc tholeiite lavas, and together with immature, terrigeneous shales or greywackes derived from erosion of the basement ttg gneisses or volcanic counterparts to them, with a likely small contribution from tholeiitic lavas. part of the succession could also have been felsic rocks derived from erosion of the basement ttg gneisses or volcanic counterparts to the latter, with a likely small contribution from tholeiitic lavas. the whole series is capped by a layer c. 100 m thick of mafic volcanic rocks likewise of tholeiitic affinity. all these lithological components are in good 50 agreement with either an arc-related setting, with a plutonic arc developing simultaneously with the filling of volcano-detritic basins with lavas of similar affinities and immature sediments; or with an accretionary wedge environment involving ocean floor juxtaposed together with similar sediments. in both cases, they correspond to an active subduction margin. intense migmatisation of the basement is probably associated with the emplacement of the anatectic, likely synkinematic kangaatsiaq granite. this was apparently synchronous with an early, lower granulite facies (d1/d1b) deformation event that may have ended with strain relaxation and exhumation of the rocks from the active margin at the end of the archaean cycle. the supracrustal association and the sequence of events in the kangaatsiaq area are comparable to the evolution of many archaean greenstone belts (e.g. card 1990; chadwick et al. 1996; hunter et al. 1998). on the other hand, classical archaean components such as orthochemical sediments and plume-related komatiites (arndt 1994) or orthochemical components (lowe 1994) are completely missing from the kangaatsiaq area. however, this apparently rather uncommon absence is known from other midto late archaean greenstones, also in west greenland (e.g. garde 1997). the setting is sometimes interpreted as being arc-related (card 1990; lowe 1994). in contrast, widespread melting and granite emplacement at the end of the archaean is a very common situation, which has been described in many studies (among others, e.g. gorman et al. 1978; card 1990; sylvester 1994; windley 1995; chadwick et al. 1996; moyen et al. 2003b). acknowledgements j.a.m. van gool, g.i. alsop, s. piazolo and s. mazur visited the area in 2001, and their work was used as a basis for the subsequent mapping. they also provided useful comments on the geology and metamorphic history of the region. a.a. garde kindly supplied analyses of the basement gneisses; his editorial help with the figures and manuscript is also gratefully acknowledged. reviews by a.g. leslie and a second reviewer greatly improved the original manuscript. linguistic corrections by r.w. belcher were also of greatest help. chemical analyses were performed at geus. references arndt, n.t. 1994: archean komatiites. in: condie, k.c. 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(ed.): andesites, 525–548. chichester: wiley. piazolo, s. 2002: overview of the metamorphic evolution of tonalitic gneisses and metasedimentary sequences from the kangaatsiaq, lersletten and sydostbugten area – first comparison to adjacent areas. in: nielsen, b.m. & thrane, k. (eds): workshop on nagssugtoqidian and rinkian geology, west greenland. danmarks og grønlands geologiske undersøgelse rapport 2002/9, 32–33. piazolo, s., alsop, g.i., møller nielsen, b. & van gool, j.a.m. 2004: the application of gis to unravel patterns of deformation in high grade terrains: a case study of indentor tectonics from west greenland. in: alsop, g.i. & holdsworth, r.e. (eds): flow processes in faults and shear zones. geological society special publication (london) 224, 63–78. roser, b.p. & korsch, r.j. 1988: provenance signature of sandstonemudstone suite determined using discriminant function analysis of major-element data. chemical geology 67, 119–139. shervais, j.w. 1982: ti-v plots and the petrogenesis of modern and ophiolitic lavas. earth and planetary science letters 59, 101–118. 52 sørensen, k., glassley, w., korstgård, j. & stensgaard, b.m. 2006: the nordre strømfjord shear zone and the arfersiorfik quartz diorite in arfersiorfik, the nagssugtoqidian orogen, west greenland. in: garde, a.a. & kalsbeek, f. (eds): precambrian crustal evolution and cretaceous–palaeogene faulting in west greenland. geological survey of denmark and greenland bulletin 11, 145–161 (this volume). steenfelt, a., garde, a.a. & moyen, j.-f. 2005: mantle wedge involvement in the petrogenesis of archaean grey gneisses in west greenland. lithos 79, 207–228. sylvester, p.j. 1994: archean granitic plutons. in: condie, k.c. (ed.): archean crustal evolution. developments in precambrian geology 11, 261–314. taylor, s.r. & mclennan, s.m. 1985: the continental crust: its composition and evolution, 312 pp. oxford: blackwell. thrane, k.t. & connelly, j. 2002: linking the nagssugtoqidian orogen and the rinkian belt: preliminary ages from the disko bugt region. in: nielsen, b.m. & thrane, k. 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(eds): precambrian crustal evolution and cretaceous–palaeogene faulting in west greenland. geological survey of denmark and greenland bulletin 11, 125–144 (this volume). van gool, j.a.m. et al. 2002a: precambrian geology of the northern nagssugtoqidian orogen, west greenland: mapping in the kangaatsiaq area. geology of greenland survey bulletin 191, 13–23. van gool, j.a.m., connelly, j.n., marker, m. & mengel, f.c. 2002b: the nagssugtoqidian orogen of west greenland: tectonic evolution and regional correlations from a west greenland perspective. canadian journal of earth sciences 39, 665–686. van kranendonk, m.j. 2003: archaean tectonics in 2001: an earth odyssey. precambrian research 127, 1–3. vernon, r.h. 2000: review of microstructural evidence of magmatic and solid-state flow. electronic geosciences 5(2). willigers, b.j.a., van gool, j.a.m., wijbrans, j.r., krogstad, e.j. & mezger, k. 2002: posttectonic cooling of the nagssugtoqidian orogen and a comparison of contrasting cooling histories in precambrian and phanerozoic orogens. journal of geology 110, 503–517. windley, b.f. 1995: the evolving continents, 3rd ed., 526 pp. chester: john wiley and sons. _______________________________________________________________________________________________________________________________________________________________________________________________ manuscript received 7 june 2004; revision accepted 1 february 2006 geological survey of denmark and greenland bulletin 4, 2003, pp 89-92 89 a major cliff collapse took place at store stejlebjerg in the southern part of møns klint on 5 july 2003 (fig. 1). this cliff collapse was one in a number of rock falls that has affected møns klint with a frequency of about one per five years. geological investigations of the rock fall at store stejlebjerg were carried out by the geological survey of denmark and greenland (geus) after the danish forest and nature agency had asked for advice and help concerning security regulations for public access to the site. geus was prepared for this type of investigation due to the survey’s engagement in the european union project protect, which aims at prediction of chalk cliff collapses. in this project a number of sites in northern europe have been selected for detailed investigation, among which two are situated at møns klint, southeast denmark (fig. 1). this report provides a short description of the 2003 cliff collapse at møns klint and a brief description of the protect project and its practical implications for cliff collapse evaluation. the store stejlebjerg cliff collapse møns klint is a 4 km long n–s-trending coastal chalk cliff up to 130 m high bordering the east side of the island of møn (fig. 1). this cliff provides an instructive structural cross-section through a large glaciotectonic complex, the southern part of which can be characterised as an imbricate fan with thrust sheets consisting of c. 60 m upper maastrichtian chalk overlain by 10−15 m glacial deposits of weichselian age (surlyk & håkansson 1999; pedersen 2000). store stejlebjerg forms a nearly 90 m high vertical cliff section that prior to the rock fall had an irregular overhang from about 25 m up to about 70 m a.s.l. (fig. 2). this overhang formed the basal boundary of the collapse unit, which is estimated to have a volume of 8000 m3. the large debris fans formed by the rock fall, constituted a southern 70 m long seawards projecting peninsula consisting of large blocks and a northern 90 m peninsula comprising finer-grained breccia (fig. 3). the site was reinvestigated in august 2003. no fractures were observed at the head of the rock fall, but fractures related to the syn-sedimentary slump deformation were seen in the lower part of the cliff with a direction parallel to the escarpment surface. slickenside surfaces on some of the largest (5 m3) blocks indicated displacements along more than one fracture plane direction. the triggering of the rock fall was interpreted to be due to a very dry spring followed by heavy rainfall immediately preceding the cliff collapse. protect the aims of the eu-project protect (prediction of the erosion of cliffed terrains), eu contract no. evk3-ct2000-00029, are to develop predictive tools that will identify sections of coastal chalk cliffs that are approaching a state of imminent collapse and allow accurate forecasts to be made concerning the timing of the collapse. the monitoring tools geological survey of denmark and greenland bulletin 4, 89–92 (2004) © geus, 2004 prediction and risk evaluation of chalk cliff collapse: the protect project stig a. schack pedersen and ingelise møller fig. 1. location map of the five test sites selected for the investigation of chalk cliff collapse in the eu-project protect. focused on are azimuthal resistivity measurements, microseismicity and acoustic emission, and the project also aims at contributing to the understanding of the physical properties of rock masses which lead to unstable cliffs and failures. the project works closely with user communities in order to ensure that development is adapted to user requirements. this includes issuing informed hazard warnings in areas around cliffs, providing information for land-use planning in the coastal zone and conservation regulations, and maximising the use of the cliffed coastline as an amenity (busby et al. 2002). in the protect project nine partners are involved: (1) british geological survey – project coordinator; (2) university of brighton, england; (3) bureau de recherche géologiques et minières, france; (4) geological survey of denmark and greenland; (5) institut national de l’environement industriel et des risques, france; (6) isle of wight centre for the coastal environment, england; (7) direction departementale de l’equipement de la seine maritime, france; (8) urzad morskiw gdyni, poland; and (9) consorzio ferrara ricerche, italy. the main scientific topics addressed by the project are discussed below. detection of fracture dilatancy temporal azimuthal apparent resistivity measurements are made at the five research sites at bimonthly intervals (figs 1, 4). the parameters required to monitor variations in the rock mass are determined and fracture orientations are calculated for each research site. detection of cracking five accelerometers and five geophones were installed at the test site in france (mesnil val; fig. 1) and acquisition began in january 2002. initial investigations of the microseismic activity shows that a microseismic event can be recorded on one transducer. as a consequence, the microseismic network is strongly recommended compared to the waveguide system, which has also been considered (busby et al. 2002). influence of rock and external parameters detailed rock mass data and erosion data are collected from the research sites, and rock sampling for investigation of physical properties and strength of the chalk is carried out. meteorological and water level data, including external temperature, barometric pressure, wind velocity and direction, and precipitation are recorded for the study of influence of external parameters on the behaviour of the rock mass. 90 fig. 2. oblique aerial photograph of the c. 90 m high chalk cliff at store stejlebjerg prior to the rock fall. the photograph was taken on 13 march 2003 as part of a systematic photogrammetric survey aiming at an aerotriangulation of the møns klint cross-section for future examinations of structural geology and landslide activity. fig. 3. the broad 90 m long peninsula formed by the rock fall seen from the top of store stejlebjerg on 29 august 2003. the preservation potential of these breccias is not high and the peninsula will probably be lost to erosion in about three years. 91 fig. 4. maps of the five test sites established for the protect investigations. all locality maps drawn with north at top. contour lines are in metres. each grid point is provided with a vector bar representing the displacement of the grid point based on repeated measurements. the displacement is relative to the first measurement. horizontal displacements shown by length of bar (upscaled by a factor of 100). vertical displacments are shown by colour. 92 interpretation and integration of data a database has been established for the data sets collected by the protect project. all surveys of fracturing and rock strength are stored for comparison and integrated interpretation. a rock fall at the french site in june 2002 destroyed all the connecting wires and conductors installed in january, and had to be replaced. however, the signal obtained from this rock fall gave a very marked precursor anomaly, which indicated that the collapse could be predicted about eight hours before it happened. during the first project period a monitoring system for direct verification of precursor movements in the terrains or actual collapse of the cliff was established. five field sites were selected for the investigations: two at møns klint in denmark (jættebrinken and dronningestolen), two sites at eastbourne on the south coast of england (beachy head and birling gap), and one site at the french coast in normandy (mesnil val; fig. 1). subsequent reporting from the monitoring system has been submitted regularly, and measurements of the test grids established in the field have been carried out at an interval of about 4 months. for testing the dislocations in the terrain a grid has been established at each research field (e.g. fig. 4). the grid consists of a number of fixed points that have been measured with a theodolite. in general the deviation of the measurements is within ± 1 cm on the z-coordinate. the x and y coordinates have a somewhat larger deviation, which is mostly below ± 3 cm (pedersen et al. 2002). one of the results from the test grid measurements is that the volume involved in the collapse at mesnil val in june 2002 could be calculated to 2750 m3. moreover, it is evident that initial creep along the fault zone displacing the birling gap field has been documented by the test grid measurements (fig. 4). prediction and risk evaluation of landslides and rock falls in order to predict rock falls the sites for potential collapse must be identified and the triggering mechanism has to be understood. the sites with most potential for collapse are of course overhanging cliffs; however, some of the collapses are also related to older fault and fracture systems, which demand a structural analysis to provide a structural model for prediction of risky sites. geus will continue its mapping of møns klint, and when the potential sites for collapse have been identified it will be possible to install monitoring equipment for collapse warning. the experiences from protect suggest that acoustic emission is a promising tool, and the survey intends to continue cooperation for developing improved prediction equipment. considering the risk evaluation, one evaluation parameter is the past frequency of rock falls. the record of landslides at møns klint over the last 100 years indicates that a major landslide or rock fall will occur about every fifth year (pedersen 2003). another parameter could be to estimate the erosion rate. at møns klint the erosion rate varies from zero at jættebrinken to nearly 50 cm per year north of dronningestolen (pedersen 2003). the average erosion rate is about 35 cm per year, which is relatively small compared to the erosion rate of about 70 cm per year along the english channel (data from the protect data files). in spring 2003, geus was asked to provide a landslide risk analysis for a projected exhibition centre at møns klint. based on a structural model for the location of the centre and the erosion rate estimates the main conclusion was that the centre would not be threatened by coastal landslide erosion in the foreseeable future (pedersen 2003). references busby, j.p., gourry, j.c., senfaute, g., pedersen, s. & mortimore, r. 2002: can we predict coastal cliff failure with remote, indirect measurements. in: mcinnes, r. & jakeways, j. (eds): instability, planning and management, 203–208. london: thomas telford. pedersen, s.a.s. 2000: superimposed deformation in glaciotectonics. bulletin of the geological society of denmark 46, 125–144. pedersen, s.a.s. 2003: vurdering af skredrisiko for området oven for maglevandsfaldet på møns klint. danmarks og grønlands geologiske undersøgelse rapport 2003/50, 35 pp. pedersen, s.a.s., møller, i. & gudmunsson, l. 2002: test grid established for the eu-project protect at cliffed terrains in denmark, england and france. danmarks og grønlands geologiske undersøgelse rapport 2002/30, 30 pp. surlyk, f. & håkansson, e. 1999: maastrichtian and danian strata in the southeastern part of the danish basin. in: pedersen, g.k. & clemmensen, l.b. (eds): field trip guide for 19th regional european meeting of sedimentology (august 1999), 29–58. copenhagen: ias. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: sasp@geus.dk << /ascii85encodepages false /allowtransparency false /autopositionepsfiles true /autorotatepages /all /binding /left /calgrayprofile (dot gain 20%) /calrgbprofile (srgb iec61966-2.1) /calcmykprofile (u.s. web coated \050swop\051 v2) /srgbprofile (srgb iec61966-2.1) /cannotembedfontpolicy /warning /compatibilitylevel 1.4 /compressobjects /tags /compresspages true /convertimagestoindexed true /passthroughjpegimages true /createjdffile false /createjobticket false /defaultrenderingintent /default /detectblends true /colorconversionstrategy /leavecolorunchanged /dothumbnails false /embedallfonts true /embedjoboptions true /dscreportinglevel 0 /syntheticboldness 1.00 /emitdscwarnings false /endpage -1 /imagememory 1048576 /lockdistillerparams 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<feff4f7f752890194e9b8a2d5b9a5efa7acb76840020005000440046002065874ef65305542b8f039ad876845f7150cf89e367905ea6ff0c4fbf65bc63d066075217537054c18cea3002005000440046002065874ef653ef4ee54f7f75280020004100630072006f0062006100740020548c002000520065006100640065007200200035002e0030002053ca66f465b07248672c4f86958b555f3002> >> >> setdistillerparams << /hwresolution [2400 2400] /pagesize [595.000 842.000] >> setpagedevice methods article | short koch & schneider 2022: geus bulletin 49. 8292. https://doi.org/10.34194/geusb.v49.8292 1 of 7 long short-term memory networks enhance rainfall-runoff modelling at the national scale of denmark julian koch* , raphael schneider department of hydrology, geological survey of denmark and greenland, copenhagen, denmark abstract this study explores the application of long short-term memory (lstm) networks to simulate runoff at the national scale of denmark using data from 301 catchments. this is the first lstm application on danish data. the results were benchmarked against the danish national water resources model (dk-model), a physically based hydrological model. the median kling-gupta efficiency (kge), a common metric to assess performance of runoff predictions (optimum of 1), increased from 0.7 (dk-model) to 0.8 (lstm) when trained against all catchments. overall, the lstm outperformed the dk-model in 80% of catchments. despite the compelling kge evaluation, the water balance closure was modelled less accurately by the lstm. the applicability of lstm networks for modelling ungauged catchments was assessed via a spatial split-sample experiment. a 20% spatial hold-out showed poorer performance of the lstm with respect to the dk model. however, after pre-training, that is, weight initialisation obtained from training against simulated data from the dk-model, the performance of the lstm was effectively improved. this formed a convincing argument supporting the knowledge-guided machine learning (ml) paradigm to integrate physically based models and ml to train robust models that generalise well. introduction the runoff at a given point along a river network can be defined as the outflow generated in the upstream contributing area. accurate modelling of runoff has been a prime research theme for several decades (wagener et al. 2004). a multitude of numerical modelling tools, from parsimonious conceptual rainfall-runoff models to complex fully distributed physically based models (pbms), have been developed. in recent years, machine learning (ml) models, in particular, long short-term memory (lstm) networks, have proved useful for rainfall-runoff modelling. since the first application by kratzert et al. (2018), lstms quickly gained popularity and have typically outperformed traditional hydrological models under data-rich settings (mai et al. 2021) and in ungauged catchments (kratzert et al. 2019a). the knowledge-guided ml paradigm aims to increase robustness and generalisability by integrating scientific knowledge into ml models (nearing et al. 2020; reichstein et al. 2019). this can be achieved by building physical constraints, such as the first-principle law of mass conservation (hoedt et al. 2021), into a ml model or using a pbm to augment training data (jia et al. *correspondence: juko@geus.dk received: 17 aug 2021 accepted: 07 dec 2021 published: 13 jan 2022 keywords: rainfall-runoff modelling, long-short term memory networks, deep learning, knowledge-guided machine learning, pre-training-finetuning abbreviations: camels: catchment attributes and meteorology for large-sample studies fbal: flow balance kge: kling-gupta efficiency lstm: long short-term memory ml: machine learning mse: mean squared error pbm: physically based model geus bulletin is an open access, peerreviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: hyojin kim (geus, denmark) reviewed by: two anonymous reviewers. funding: none provided competing interests: see page 6 additional files: see page 6 https://doi.org/10.34194/geusb.v49.8292 https://orcid.org/0000-0002-7732-3436 https://orcid.org/0000-0001-9628-0809 mailto:juko@geus.dk koch & schneider 2022: geus bulletin 49. 8292. https://doi.org/10.34194/geusb.v49.8292 2 of 7 www.geusbul let in.org 2021). in this context, the method of pre-training by weight initialisation using pbm simulation data appears to be very promising, as a pre-trained lstm attempts to emulate a pbm. the rapid advancement of ml models for runoff prediction was facilitated by the availability of multiple large-scale runoff data sets containing a long timeseries of observed runoff, dynamic meteorological forcing and static catchment attributes, referred to as catchment attributes and meteorology for large-sample studies (camels) data sets (e.g. addor et al. 2017). in this article, we highlight the value of danish hydrological big data for the advancement of ml-based runoff modelling. the danish case offers a data-rich setting with over 300 stations and high-quality auxiliary data. moreover, there exists a national water resources model (the dk-model), an advanced hydrological pbm that integrates groundwater and surface water processes (højberg et al. 2013; stisen et al. 2019). the dk-model is a perfect benchmark for ml model development and provides simulated runoff, which is valuable for augmentation, as well as auxiliary hydrological information, such as groundwater conditions. in this study, we aim to (1) highlight the value of danish hydrological big data for advancing ml research at an international level, (2) implement a state-of-the-art lstm to model runoff at the national scale of denmark and (3) test a knowledge-guided lstm based upon pre-training against simulated runoff obtained from a pbm. methods data as in existing camels data sets, we curated a data set comprising observed runoff as well as dynamic and static attributes for 301 danish catchments (fig.  1). the catchments vary in size between 10 km2 and 2574 km2 with an average of 133 km2. the dynamic variables cover a period of 21 years (1990–2011) at daily timesteps and comprise observed runoff, simulated runoff (dk-model), air temperature, potential evapotranspiration and precipitation (fig. 2). the three meteorological variables were derived from gridded data provided by the danish meteorological institute and represent daily-averaged conditions for the entire catchment (scharling 1999a, 1999b). a complete timeseries of 21 years of daily observed runoff were available for 51% of the catchments, with 77% of the catchments having at least an 80% coverage. the runoff was normalised by the catchment size to mm/ day to give equal weight to the catchments during training, independent of their size. in total, 17 static catchment attributes were compiled. eleven of which were calculated as catchment averages: precipitation, potential evapotranspiration, air temperature, slope, topographic wetness index, clay fraction, annual, summer and winter-simulated water table depth (dk-model), exceedance probability of a simulated water-table depth less than 1 m (dk-model) and the thickness of the surficial clay layer. five land-use classes were expressed as percentages: forest, wetland, lake, agriculture and urban. finally, the catchment area was included as well. all data are available at https://doi.org/10.22008/fk2/ycqxtr. long short-term memory the lstm network architecture is a special type of recurrent neural network, designed to store and regulate information over time, which makes lstms well suited to learn long-term dependencies and memory effects (hochreiter & schmidhuber 1997). the lstm is described in full elsewhere (kratzert et al. 2018; shen 2018). similar to traditional hydrological models, the lstm processes input data time step after time step. runoff on a specific day is simulated based on the timeseries of length n of the preceding n days of meteorological data. kratzert at al. (2019b) developed the entity-aware lstm, which is an adaptation of the standard lstm capable of learning catchment similarities based on the static attributes, which are treated in a separate embedding layer. in this study, we applied the proposed entity-aware lstm, referred to simply as lstm hereafter. we used the fig. 1 map of denmark showing the 301 catchments used in this study. 60 catchments were randomly sampled for the spatio-temporal split-sample experiment. ! ! ! ! ! ! ! ! ! ! ! ! ! !! ! ! ! ! !! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! !! ! ! ! ! ! ! ! ! ! ! ! ! ! !! ! !! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! !! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! !! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! !! ! ! ! ! !! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! ! 60 800 20 40 km q stations spatial split-sample catchments https://doi.org/10.34194/geusb.v49.8292 http://www.geusbulletin.org https://doi.org/10.22008/fk2/ycqxtr koch & schneider 2022: geus bulletin 49. 8292. https://doi.org/10.34194/geusb.v49.8292 3 of 7 www.geusbul let in.org neuralhydrology codebase (github.com/neuralhydrology/neuralhydrology/) to train and evaluate the models used in this study. experimental setup hyperparameters and general settings as the purpose of this study was to initially explore the lstm applicability, hyperparameters were not optimised. following kratzert et al. (2019b), we assigned the following hyperparameter values: a learning rate of 0.001, a batch size of 256, an input length of 270 days, 64 hidden cell states, a dropout rate of 0.4 and 20 training epochs. all models were trained with five different seeds and the average of the five models was used for the final lstm prediction. the model setup files are available at https://doi.org/10.22008/ fk2/ycqxtr. split-sample experiments we conducted both a temporal split-sample and a spatio-temporal split-sample experiment to test the capabilities of a lstm for danish runoff data. the temporal split-sample experiment used data from all 301 stations for training. the timeseries were split into a training period of 11 years (2000–2011) and a test period of 10 years (1990–1999; fig. 2). the two periods correspond to the calibration and test period of the dk-model, which permitted a fair comparison between the two models. the spatio-temporal split-sample experiment was divided into the same training and test periods. furthermore, 20% of the stations were randomly selected and removed from the training data set and retained for model evaluation of the spatio-temporal split-sample experiment (i.e. a 20% spatial hold-out; fig. 1). this experiment offers a more robust evaluation, as it tests the transferability of 80% of stations to the remaining 20%. this allows us to assess the ability to predict ungauged basins. pre-training the concept of pre-training can be used to initialise the weights of a lstm using alternative runoff data before fine-tuning the lstm using the actual runoff data from the catchments of interest. runoff data for pre-training can potentially be obtained from observational data sets from a larger or different geographical region or from a pbm. in this study, we followed the latter and employed simulation data from the dk-model to pretrain. in this way, the lstm aimed to emulate the process descriptions of the pbm before being fine-tuned against observed runoff. the training epochs were set to 15 for pre-training and 5 for fine-tuning. simulated runoff at all 301 stations for the training period of 11 years (2000–2011) was used for pre-training, and it was applied to both split-sample experiments. evaluation metrics for training the lstm network, the mean squared error (mse) between the observed and simulated runoff was selected as the loss function. two alternative metrics were calculated for the model evaluation, namely the kling-gupta efficiency (kge) and the averaged flow balance (fbal). kge is a three-component metric that considers the correlation, the standard deviation ratio and the bias between the observed and simulated runoff (gupta et al. 2009). fbal quantifies the water balance closure between the observed and simulated runoff relative to the observed flow (henriksen et al. 2003). negative fbal scores indicate an overestimation of the model with respect to the observed runoff. the optimal values for kge and fbal are 1 and 0, respectively. results and discussions the cumulative density functions for kge and fbal are presented in figure 3. the lstm was benchmarked against the dk-model (pbm), and the effect of pre-training was also investigated. superior performance could be attributed to the lstm, with and without pre-training, fig. 2 dynamic input data for a single catchment used to train the lstm. a: precipitation. b: potential evapotranspiration. c: air temperature. d: observed runoff (obs) and simulated runoff (pbm) were used as training data. the training period and test period are shown in a. b: potential evapotranspiration c: air temperature d: runoff a: precipitation test training pbm obs m 3 /s m m /d ay m m /d ay 0 10 20 19 92 19 94 19 96 19 98 20 00 20 02 20 04 20 06 20 08 20 10 19 90 40 ° c 20 20 0 0 0 2 4 https://doi.org/10.34194/geusb.v49.8292 http://www.geusbulletin.org http://github.com/neuralhydrology/neuralhydrology/ http://github.com/neuralhydrology/neuralhydrology/ https://doi.org/10.22008/fk2/ycqxtr https://doi.org/10.22008/fk2/ycqxtr koch & schneider 2022: geus bulletin 49. 8292. https://doi.org/10.34194/geusb.v49.8292 4 of 7 www.geusbul let in.org with respect to kge for the temporal split-sample experiment. the median kge was 0.8 for both lstm configurations and 0.7 for the pbm. the conclusion was less clear for the water balance closure (fbal); here, the pbm showed normally distributed underand over-estimates with a median close to zero. however, the lstms were skewed towards negative values, that is, overestimation of runoff, with a median of –0.08. overestimated runoff was predominately evident during the low-flow summer periods. using alternative loss functions instead of the mse in the lstm training may alleviate this problem. the spatio-temporal split-sample experiment revealed that the lstm did not generalise well to ungauged basins. the median kge was 0.69 and comparable with the pbm (kge = 0.73), despite poor kge scores for the lowest 20% of the cumulative density function. the same was evident for fbal, where the lowest 20% performed poorly with respect to the pbm. however, pre-training using pbm data resulted in better performance for ungauged basins, making them comparable with the pbm. this emphasised the merit of pre-training. for the spatio-temporal split-sample experiment, where information was evidently missing in the training data set, pre-training against pbm data helped to increase performance. however, the performance did not change for the temporal split-sample experiment, where the observations provided enough information. considering kge for the temporal split-sample experiment, the lstm outperformed the pbm in 80% of catchments. this fell to 44% for the spatio-temporal experiment but could rise to 54% through pre-training. considering the absolute fbal, 68% of the catchments were simulated more precisely by the lstm than by the pbm for the temporal split-sample experiment. for the spatio-temporal split-sample experiment, this could be raised slightly from 44% to 49% through pre-training. the simulation results for two selected catchments for the 10-year test period of the temporal and spatio-temporal split-sample experiments are presented in figure 4. in the first catchment (260080), the performance between the lstms and the pbm was very comparable with a kge score of 0.8 (pbm) and 0.78 (lstm) for the temporal split-sample experiment. the performance dropped to 0.66 in the spatio-temporal split-sample experiment but increased to 0.8 through pre-training. the second catchment (420022) showed a very poor performance for the spatio-temporal split-sample experiment (kge = –0.12). however, kge improved to 0.78 through pre-training and thus became comparable with the pbm (kge = 0.75). in other words, the spatio-temporal split-sample experiment for catchment 260 080 could be simulated accurately without pre-training, because the lstm could learn the runoff behaviour of that catchment using data from similar neighbouring or upstream catchments. however, the runoff behaviour of catchment 420022 could not be learned without data from the same catchment. nevertheless, pre-training using pbm data helped to increase the performance of the lstm. figure 4 presents results for a large (260080, 323 km2) and a small (420022, 44  km2) catchment. fig. 3 cumulative density functions for kge and fbal in the test period for runoff simulated by the dk-model (pbm), the lstm model and the pre-trained lstm model (prtrn lstm). the temporal split-sample experiment is depicted in the left panels and the spatio-temporal split-sample experiment in the right panels. the optimal value of fbal is highlighted with a dashed horizontal line. the lstm predictions are based on the mean of 5 seeds, indicated here with transparent coloured lines. spatio-temporal split-sampletemporal split-sample probability probability lstm pbm fb al k g e prtrn lstm –0.5 0.0 0.5 –0.5 0.0 0.5 1.0 0.0 0.2 0.4 0.6 0.8 1.0 0.0 0.2 0.4 0.6 0.8 1.0 https://doi.org/10.34194/geusb.v49.8292 http://www.geusbulletin.org koch & schneider 2022: geus bulletin 49. 8292. https://doi.org/10.34194/geusb.v49.8292 5 of 7 www.geusbul let in.org smaller catchments generally perform less well due to the stronger imprint of anthropogenic activities (drainage and abstraction) and an increased uncertainty of precipitation data for smaller catchments. the superior performance of lstms over conceptual rainfall-runoff models or hydrological pbms was demonstrated for temporal split-sample experiments by kratzert et al. (2018), gauch et al. (2021), mai et al. (2021) and others; however, conclusions of the spatial transferability to ungauged basins are disputed. kratzert et al. (2019a) reported a superior performance of lstm for a small spatial hold-out (8%), whereas mai et al. (2021) found a worse performance for a more systematic spatial hold-out. loss function plots are presented in supplementary file s1 to elucidate the training of the applied modelling experiments in more detail. the data generally support the chosen hyperparameter values and number of training epochs. to our knowledge, this is the first study that demonstrates the merits of pre-training against pbm simulation data for runoff modelling in the context of knowledge-guided ml. in a related study, pre-training using pbm data was found to be beneficial for the modelling of lake-water temperature (read et al. 2019). for rainfall-runoff modelling, pre-training has so far been found to be suitable for transferring trained lstms from one geographical region to another (ma et al. 2021). we have shown that pre-training using pbm data offers great potential to initialise the lstm with diverse runoff behaviour. here, we constrained only the pre-training to the same catchments and time; however, in theory, pbm simulations for different climate change scenarios or a larger geographical domain could inform the lstm with diverse runoff behaviour not seen in the observed runoff data. most of the published studies on lstm runoff modelling are of catchments with a low anthropogenic fig. 4 two example catchments showing the observed (obs) and simulated runoff for the test period. a: catchment 260080, 323 km2. b: catchment 420022, 44 km2. simulated data comprise the dk-model (pbm), the lstm-based model and the pre-trained lstm-based model (prtrn lstm). the runoff predictions of the lstm-based models are given for the temporal split-sample (ts) and spatio-temporal split-sample (sts) experiments. a: runoff 260080 b: runoff 420022 m 3 /s m 3 /s 15 10 20 25 5 0 1.5 1.0 2.0 0.5 0.0 1990 1991 1992 1993 1994 1995 1996 1997 1998 1999 2000 lstm_ts lstm_sts prtrn lstm_ts prtrn lstm_sts pbm obs https://doi.org/10.34194/geusb.v49.8292 http://www.geusbulletin.org koch & schneider 2022: geus bulletin 49. 8292. https://doi.org/10.34194/geusb.v49.8292 6 of 7 www.geusbul let in.org impact; however, recent efforts to model highly managed catchments have documented promising results as well (ouyang et al. 2021). the 301 danish catchments selected in this study are, to a large degree, affected by groundwater abstraction and drainage, and the effect of the degree of anthropogenic impact on model performance and transferability should be investigated in future work. conclusions we draw the following main conclusions from the initial application of lstm networks for rainfall runoff modelling at the national scale of denmark: danish hydrological big data have the potential for conducting ml research at an international level. the dk-model serves as a valuable benchmark as well as a source for augmented training data and input data in the form of static catchment attributes. an lstm can outperform a state-of-the-art hydrological model; however, accuracy decreases for ungauged catchments. this can be alleviated by pre-training against physically based simulated runoff, providing crucial information to the lstm where needed. future research studies should (1) advance knowledge-guided ml to use hydrological knowledge provided by the dk-model optimally; (2) test alternative lstm architectures, hyperparameters and loss functions; (3) study the effect of anthropogenic impact (drainage and groundwater abstraction) on the lstm; (4) investigate ways of interpreting lstm models to gain new insights into the runoff process in denmark; (5) apply a broad range of hydrological signatures in the evaluation of lstms; and (6) produce a camels data set for denmark to provide high-quality hydrological and meteorological data. acknowledgements the authors acknowledge the developer team behind the neuralhydrology codebase for making lstm modelling tools so accessible. furthermore, geus colleagues h.j. henriksen and s. stisen are thanked for providing valuable feedback to this manuscript. two anonymous reviewers are thanked for providing valuable comments to this manuscript. author contributions jk: code development, writing original draft and visualisation. rs: data  preparation. both authors have  conceptualised the study and  design, read, edited and agreed to the published version of the manuscript. competing interests the authors declare no competing interests. additional files all data and model setups are available at: https://doi.org/10.22008/ fk2/ycqxtr. an additional supplementary file is available at: https:// doi.org/10.22008/fk2/wcf76i. references addor, n., newman, a.j., mizukami, n. & clark, m.p. 2017: the camels data set: catchment attributes and meteorology for large-sample studies. hydrology and earth system sciences 21, 5293–5313. https:// doi.org/10.5194/hess-21-5293-2017 gauch, m., mai, j. & lin, j. 2021: the proper care and feeding of camels: how limited training data affects streamflow prediction. environmental modelling and software 135, 104926. https://doi.org/10.1016/j. envsoft.2020.104926 gupta, h.v, kling, h., yilmaz, k.k. & martinez, g.f. 2009: decomposition of the mean squared error and nse performance criteria: implications for improving hydrological modelling. journal of hydrology 377(1–2), 80–91. https://doi.org/10.1016/j.jhydrol.2009.08.003 henriksen, h.j., troldborg, l., nyegaard, p., sonnenborg, t.o., refsgaard, j.c. & madsen, b. 2003: methodology for construction, calibration and validation of a national hydrological model for denmark. journal of hydrology 280, 52–71. https://doi.org/10.1016/s0022-1694(03)00186-0 hochreiter, s. & schmidhuber, j. 1997: long short-term memory. neural computation 9, 1735–1780. https://doi.org/10.1162/neco.1997.9.8.1735 hoedt, p.-j., kratzert, f., klotz, d., halmich, c., holzleitner, m., nearing, g., hochreiter, s. & klambauer, g. 2021: mc-lstm: mass-conserving lstm. arxiv preprint arxiv:2101.05186 (2021). højberg, a.l., troldborg, l., stisen, s., christensen, b.b.s. & henriksen, h.j. 2013: stakeholder driven update and improvement of a national water resources model. environmental modelling and software 40, 202–213. https://doi.org/10.1016/j.envsoft.2012.09.010 jia, x. et al. 2021: physics-guided recurrent graph model for predicting flow and temperature in river networks. in: demeniconi, c. & davidson, i. 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water temperature. water resources research 55, 9173–9190. https://doi.org/10.1029/2019wr024922 reichstein, m., camps-valls, g., stevens, b., jung, m., denzler, j., carvalhais, n. & prabhat 2019: deep learning and process understanding for data-driven earth system science. nature 566(7743), 195–204. https://doi.org/10.1038/s41586-019-0912-1 scharling, m. 1999a: klimagrid danmark: nedbør, lufttemperatur og potentiel fordampning 20*20 & 40*40 km. danish meteorological institute technical report 99-12, dmi, copenhagen, dk. https://doi.org/10.34194/geusb.v49.8292 http://www.geusbulletin.org https://doi.org/10.22008/fk2/ycqxtr https://doi.org/10.22008/fk2/ycqxtr https://doi.org/10.22008/fk2/wcf76i https://doi.org/10.22008/fk2/wcf76i https://doi.org/10.5194/hess-21-5293-2017 https://doi.org/10.5194/hess-21-5293-2017 https://doi.org/10.1016/j.envsoft.2020.104926 https://doi.org/10.1016/j.envsoft.2020.104926 https://doi.org/10.1016/j.jhydrol.2009.08.003 https://doi.org/10.1016/s0022-1694(03)00186-0 https://doi.org/10.1162/neco.1997.9.8.1735 https://doi.org/10.1016/j.envsoft.2012.09.010 https://doi.org/10.1137/1.9781611976700.69 https://doi.org/10.1137/1.9781611976700.69 https://doi.org/10.5194/hess-22-6005-2018 https://doi.org/10.5194/hess-22-6005-2018 https://doi.org/10.1029/2019wr026065 https://doi.org/10.5194/hess-23-5089-2019 https://doi.org/10.5194/hess-23-5089-2019 https://doi.org/10.1029/2020wr028600 https://doi.org/10.1061/(asce)he.1943-5584.0002097 https://doi.org/10.1061/(asce)he.1943-5584.0002097 https://doi.org/10.1029/2020wr028091 https://doi.org/10.1016/j.jhydrol.2021.126455 https://doi.org/10.1029/2019wr024922 https://doi.org/10.1038/s41586-019-0912-1 koch & schneider 2022: geus bulletin 49. 8292. https://doi.org/10.34194/geusb.v49.8292 7 of 7 www.geusbul let in.org scharling, m. 1999b: klimagrid danmark: nedbør 10*10 km (ver. 2). danish  meteorological institute technical report 99-15, dmi, copenhagen, dk. shen, c. 2018: a trans-disciplinary review of deep learning research and  its relevance for water resources scientists. water resources research 54, 8558–8593. https://doi. org/10.1029/2018wr022643 stisen, s., ondracek, m., troldborg, l., schneider, r.j.m. & van thil, m.j. 2019: national vandressource model. modelopstilling og kalibrering af dk-model 2019. danmarks og grønlands geologiske undersøgelse rapport 2019/31, geus, copenhagen, dk. wagener, t., wheater, h.s. & gupta, h.v. 2004: rainfall-runoff modelling in gauged and ungauged catchments, 332 pp. london: imperial college press. https://doi.org/10.1142/p335 https://doi.org/10.34194/geusb.v49.8292 http://www.geusbulletin.org https://doi.org/10.1029/2018wr022643 https://doi.org/10.1029/2018wr022643 https://doi.org/10.1142/p335 long short-term memory networks enhance rainfall-runoff modelling at the national scale of denmark abstract acknowledgements references figures fig. 1 map of denmark showing the 301 catchments used in this study. 60 catchments were randomly sam fig. 2 dynamic input data for a single catchment used to train the lstm. a: precipitation. b: potent fig. 3 cumulative density functions for kge and fbal in the test period for runoff simulated by the fig. 4 two example catchments showing the observed (obs) and simulated runoff for the test period. a method article | short falk & madsen 2023: geus bulletin 53. 8353. https://doi.org/10.34194/geusb.v53.8353 1 of 7 machine learning-based estimation and clustering of statistics within stratigraphic models as exemplified in denmark frederik alexander falk*1 , rasmus bødker madsen2 1department of geoscience, aarhus university, aarhus, denmark; 2department of near surface land and marine geology, geological survey of denmark and greenland (geus), aarhus, denmark abstract estimating a covariance model for kriging purposes is traditionally done using semivariogram analyses, where an empirical semivariogram is calculated, and a chosen semivariogram model, usually defined by a sill and a range, is fitted. we demonstrate that a convolutional neural network can estimate such a semivariogram model with comparable accuracy and precision by training it to recognise the relationship between realisations of gaussian random fields and the sill and range values that define it, for a gaussian type semivariance model. we do this by training the network with synthetic data consisting of many such realisations with the sill and range as the target variables. because training takes time, the method is best suited for cases where many models need to be estimated since the actual estimation itself is about 70 times faster with the neural network than with the traditional approach. we demonstrate the viability of the method in three ways: (1) we test the model’s performance on the validation data, (2) we do a test where we compare the model to the traditional approach and (3) we show an example of an actual application of the method using the danish national hydrostratigraphic model. *correspondence: frederikfalk@geo.au.dk received: 31 may 2023 revised: 23 aug 2023 accepted: 26 sep 2023 published: 17 nov 2023 keywords: convolutional neural network, covariance model, semivariogram modelling, machine learning, local stationarity abbreviations cnn: convolutional neural network ml: machine learning nn: neural network geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: julian koch (geus, copenhagen, denmark) reviewed by: jacob skauvold (norwegian computing center, norway) funding: see page 7 competing interests: see page 7 additional files: none tabular abstract geographical coverage fyn, denmark temporal coverage n/a subject(s) covered geophysics, computational geoscience, informatics and remote sensing method type a new machine learning-based method for estimating locally optimised semivariogram parameters for grid cells in stratigraphic models followed by clustering for the introduction of the assumption of local stationarity. method name machine learning-based semivariogram model estimation and clustering instruments and equipment used equipment used:  - a sufficiently effective computer  - matlab® software license    • machine learning toolbox for matlab    • sippi geostatistics toolbox for matlab    • mgstat geostatistics toolbox for matlab related publications none potential application(s) for this method this method may be used to infer a statistical model from one stratigraphic model, which is useful for uncertainty quantification. the method is also useful for very fast semivariogram modelling whenever the advantage of doing so outweighs the time it takes to train the model. https://doi.org/10.34194/geusb.v53.8353 https://orcid.org/0009-0003-1023-5161 https://orcid.org/0000-0001-8538-7491 mailto:frederikfalk@geo.au.dk https://creativecommons.org/licenses/by/4.0/deed.ast falk & madsen 2023: geus bulletin 53. 8353. https://doi.org/10.34194/geusb.v53.8353 2 of 7 www.geusbul let in.org introduction the properties of the subsurface are highly non-stationary, that is, they vary significantly depending on the location. consequently, in statistical descriptions of the subsurface, there is a need for mapping non-stationarity in the statistical properties and for practical purposes also defining regions wherein local stationarity can be assumed (boisvert et al. 2009). estimating the statistical properties of spatially distributed data is conventionally done through semivariogram analysis, where an experimental semivariogram is calculated as half the average squared difference between points separated by some distance, h (matheron 1963). a model is fitted to the semivariogram, which is typically defined by two parameters – a range and a sill, and oftentimes also a nugget effect (cressie 1993). alternatively, it is possible to estimate a semivariogram model by determining the maximum-likelihood combination of sill and range given the type of model. the likelihood for a semivariogram model is obtained by populating a covariance matrix using the model and taking the probability density of the corresponding multivariate normal distribution at the point defined by the data (mardia 1990). fitting a semivariogram and estimating the maximum-likelihood model share the drawbacks of being computationally intensive when many models need to be estimated. we propose a new method of estimating the sill and range given a set of spatially distributed data using machine learning (ml), specifically a convolutional neural network (cnn), which is more efficient when estimating many models. a cnn is trained to recognise the approximate mapping from a realisation of a gaussian random field to the semivariogram model that defines its probability density function. this mapping is not bijective in nature, and as such, it is not a function. however, the network should still be able to approximate a function that resembles the mapping to some degree. the idea of using a cnn for semivariogram modelling has been proposed before. one study used separate networks for interpolation and parameter estimation (jo & pyrcz 2022). others skipped parameterised models and directly estimated semivariograms for kriging (li et  al. 2022). we have chosen to focus on estimating gaussian model parameters using one network to infer their spatial distribution within large models and make the process computationally feasible. we test the method on the danish national hydrostratigraphic model (dk-model; stisen et al. 2020). we show that where local stationarity may be assumed, we can define regions within the hydrostratigraphic model with reasonable accuracy by clustering the models. required resources the required resources are as follows: • a sufficiently effective computer •  matlab® 2022b or newer – older versions have not been tested for this method. also, the machine learning toolbox for matlab, sippi geostatistics toolbox for matlab and the mgstat geostatistics toolbox for matlab •  data in the form of scattered points with a value for each point, either with irregular spacing or as a regular grid methodological protocols we use the cnn’s ability to efficiently detect structural patterns in an image, and as such, it needs a regular grid as input. we consider two cases: one where data constitute a full grid and one with scattered point data interpolated to produce a grid. for the network input, we chose the grid size 31 × 31 cells with a cell size of 100 m, and we adapted the neural network (nn) architecture from the squeezenet convolutional neural network (iandola et al. 2016), which is a native architecture in the matlab machine learning toolbox. producing training, validation and test data we produced synthetic data for the nn using the sippi toolbox in matlab (hansen et al. 2013) by taking 150 000 values for sill and range from uniform distributions, such that the sills vary between 0 m2 and 1100 m2, and the ranges vary between 100 m and 3000 m. we then simulate a realisation from each of the gaussian random fields that have the gaussian semivariance functions defined by the pairs of sill and range values. each realisation is on the 31 × 31 grid with a cell size of 100 m. to include some component of noise, we simulated and added one more realisation to each existing realisation, with the same effective range, and the sill being random between 0 m2 and the sill of the original realisation itself. figure 1 shows nine examples of the synthetic data. the synthetic data are saved both as full grids and sets of 120 points with an x-coordinate, a y-coordinate and a z-coordinate. we split the synthetic data into a training set, consisting of 90% of the data, as well as a validation and a training set, each being 5% of the data. training the network the network is trained with the ‘adam’ optimisation algorithm, and the loss function is represented by the mean squared error. we used a constant learning rate of 0.0005 and a batch size of 1000 whilst training the model https://doi.org/10.34194/geusb.v53.8353 http://www.geusbulletin.org falk & madsen 2023: geus bulletin 53. 8353. https://doi.org/10.34194/geusb.v53.8353 3 of 7 www.geusbul let in.org for 100 epochs. the total training time for the network is about 1 h. we also found that model training time could be reduced to only a few minutes if we included the fast fourier transform as a second image input channel. ultimately, we chose not to do this because that requires us to calculate the fourier transform of the data input every time we want to use the model. this is a disadvantage because it slows down the prediction process, which becomes an issue when applied to very large models. the fourier transform could be useful if training data are scarce, which could be the case with non-synthetic training data. at this point, the network is ready to use, and only requires a 31 × 31 grid input. the training can also be done with point data. however, as the cnn is based on image convolution, one should choose an interpolation method and convert the point data to a grid. validation we demonstrate the advantage of using ml for the estimation of the semivariogram model with a synthetic example, where the method is compared to the classic method of fitting an experimental semivariogram. we then demonstrate the advantage of applying the method given an actual use case. the 5% of the data, which we set aside for validation, are fed to the network and serve as an initial validation test. the size of the validation set is 7500 data points. figure 2 shows the distribution of true to predicted values for the range and sill for the validation data set. a common way to estimate the sill is to simply take the sample variance of the scattered points, which is also shown in fig. 2. the predicted range values are very close to the true values for both scattered data and full grids. the predictions using full grids are a little more accurate, which we expected since the scattered points are drawn from the full grids and thus contain less information. the predictions for the sill are equally precise between the two cases and even share the same apparent biases. for example, at sill values of 500 to 1500, the ml model over-estimates, whilst it underestimates for sill values above 2000. although we did not use bias description, it is possible by fitting a suitable polynomial function between prediction and actual values. correcting the estimates to account for bias is then straightforward. alternatively, it is reasonable to assume that improvements in the nn itself could eliminate the bias. the variance estimate is less precise but has no bias. validation on synthetic data we take a set of 1000 new synthetic realisations of size 31 × 31, each with 120 randomly drawn points, and we use these points to: 1.   perform a traditional semivariogram analysis with the following steps: a. calculate an empirical semivariogram from the points. b. fit a gaussian semivariance model to the semivariogram using a weighted leastsquares approach, where points closer to the origin have greater weight. 2.  predict the sill and range directly with our cnn with the following steps: a.  interpolate the 120 points onto the entire grid. b. pass the interpolated grid through the cnn. we then compare the accuracy of the estimates with these two methods as well as the time it takes to complete. the result of the accuracy comparison is shown in fig. 3 for eight different realisations. depending on the specific realisation, it varies whether the fitted model (black line) or cnn (blue line) is better at resolving the true model (red line). across all 1000 realisations, we see that the cnn is slightly better at approximating large values for the range than the traditional approach, whereas the traditional method is slightly better at low values. in general, both methods have similar performance. however, the time that the methods need to reach the predictions is not the same. during the test, we fig. 1 synthetic training data produced using gaussian semivariance models with random sill and range, with a component of noise. the circles show 120 randomly drawn points from each realisation. shading: smallest values are shown in blue and largest values are shown in yellow, with in-between values shown in green. sill: 455.4 range: 2713 sill: 499.1 range: 1577 sill: 1054 range: 1900 sill: 431 range: 2327 sill: 389.5 range: 1374 sill: 1080 range: 411.2 sill: 187 range: 1002 sill: 575.7 range: 2446 sill: 206.6 range: 455.1 https://doi.org/10.34194/geusb.v53.8353 http://www.geusbulletin.org falk & madsen 2023: geus bulletin 53. 8353. https://doi.org/10.34194/geusb.v53.8353 4 of 7 www.geusbul let in.org recorded the time spent for each approach and calculated the average time spent per model. the cnn approach based on existing 31 × 31 grids was able to estimate 130 semivariogram models per second, whilst for scattered data, 111 models could be estimated per second. meanwhile, the traditional semivariogram analysis, where a model is fitted to an experimental semivariogram, could only estimate 1.72 models per second. of course, these numbers depend on the computational resources available, but the important point is the relative difference in computation time between the methods. the time consumption of the ml approach is only slightly larger when interpolating point data, suggesting that the cnn consumes most of the time and not the interpolation itself. meanwhile, the traditional approach of fitting takes about 60 times more computation time. the time consumption varies quite a bit for the traditional semivariogram analysis approach, depending on the number of data points used. by using only 60 points, it may complete as many as 30 models per second; however, this is still about one-fourth of the speed of the cnn and with a significant loss of accuracy. furthermore, the cnn may also be optimised to become more efficient. validation by application of the method besides the validation on synthetic data, we also validate our method by demonstrating how it can be used to solve a real problem of obtaining non-stationary statistical properties. when dealing with models with non-stationary statistics, such as very large models like the dk-model (stisen et al. 2020), using a single semivariogram for kriging does not usually produce realistic geological structures. in our validation example, we employ the ml approach to estimate local values for the range and the sill and then use a clustering algorithm to divide the model into local regions with similar statistical properties. the algorithm is a type of unsupervised classification algorithm known as the kohonen self-organising map (kohonen 1991), which is featured as a built-in function in matlab. with this approach, the ml algorithm enables kriging with a locally optimised semivariogram model, which should be better at handling non-stationary models than approaches with a fixed semivariogram model. for the validation example, we employ this approach to the first layer in the hydrostratigrahic model on the danish island of fyn, covering roughly 3100 km2. fig. 2 the panels show the precision and accuracy of the ml model in predicting range and sill. red dots show values predicted using the scattered points and blue dots using full grids, and green dots show the estimation of sill using the sample variance from the 120 scattered points. the black trend lines indicate where points are in exact agreement with the true values. the colour-shaded regions highlight the interval within which 95% of predictions are made, bounded by the 2.5 percentile (lowermost coloured lines), 50 percentile (median; middle coloured lines) and 97.5 percentile (uppermost coloured lines). https://doi.org/10.34194/geusb.v53.8353 http://www.geusbulletin.org falk & madsen 2023: geus bulletin 53. 8353. https://doi.org/10.34194/geusb.v53.8353 5 of 7 www.geusbul let in.org the sample variance is used to estimate the sill in this example because an unbiased estimation is prioritised over a precise estimation for the purpose of clustering. as such, the example focuses on the predicted ranges and the final clustering. the top left panel in fig. 3 shows the ranges predicted with a 31 × 31 sliding window, whilst the top right panel shows the subsequent clustering result. figure 4 illustrates how the ml algorithm identifies areas of low range where the layer has more high-frequency variation and areas of high range where the layer resembles more of a smooth curve. of course, this should be seen in the context of the chosen window size, which is about 3100 m, and larger ranges than this cannot be resolved. this limitation should not pose a problem since any subsequent kriging and simulation will be modelling the residual around the sliding mean from the layer using the same window. the example layer shown here contains 497  369 individual grid cells, which means that the machine learning algorithm can complete the semivariance model estimation in about 1 h and 15 min, given a computation rate of 111 models per second. meanwhile, the traditional semivariogram analysis takes about 80 h to do the same, given a rate of 1.72 models per second. discussion and outlook in a subsurface model with a large spatial extent, the assumption of stationarity in the subsurface properties breaks down. to do proper geostatistical modelling in such a case, non-stationarity must be considered (higdon et al. 2022). non-stationarity can be modelled by introducing locally varying anisotropy (e.g. boisvert & deutsch 2011; bongajum et al. 2013; pereira et al. 2023), but these methods can be computationally challenging for large models. thus, practical tools needed for estimating the non-stationarity are currently sparse and not easily deployed for practitioners (madsen et al. 2020). here, we briefly presented a computationally efficient ml-based method that can infer gaussian properties from a stratigraphic layer model, adding a new tool to the geostatistician’s toolbox to solve issues of non-stationarity. during testing, several different ml approaches were tried, including regression trees, random forest and a classical nn, but the deployment and adoption of a cnn were the most successful. it is known that neural networks in general are universal approximators that can approximate any continuous lebesgue integrable function. this was proven for networks with a fixed number of hidden layers and an arbitrary number of neurons, also known as the arbitrary width case (hornik 1991). recently, it was also proven for relu nns with a fixed synthetic example 10 20 30 y [m ] semivariance [m2] synthetic example semivariance [m2] 10 20 30 y [m ] 10 20 30 y [m ] 10 20 30 x [m] 10 20 30 y [m ] 0 1000 2000 3000 distance [m] 10 20 30 x [m] 0 1000 2000 3000 distance [m] synthetic data data true model ml model fitted model fig. 3 synthetic examples: 8 realisations are shown – each of their own gaussian semivariogram model with some component of noise. the semivariance models show a high level of accuracy for the traditional fitting approach as well as the ml approach. https://doi.org/10.34194/geusb.v53.8353 http://www.geusbulletin.org falk & madsen 2023: geus bulletin 53. 8353. https://doi.org/10.34194/geusb.v53.8353 6 of 7 www.geusbul let in.org number of neurons (fixed width) and an arbitrary number of hidden layers, also known as the arbitrary depth case (zhou et al. 2017). in the presented use-case, the cnn probably produced better estimates compared to the other ml approaches because the convolution of different layers makes the cnn better at analysing the spatial information in the training data. with the rapid development in ml algorithms, the cnn might soon be outperformed; however, the methodology of training the ml model to recognise gaussian covariances from point data does not change but can only improve its precision with new algorithms. this presents an improvement over, for example, a traditional semivariogram analysis, the performance of which comes with a tradeoff between the ability to accurately predict shorter versus longer ranges. this trade-off occurs due to the constraints imposed by the weights on the experimental variogram at different range intervals. the nn architecture used from squeezenet (iandola et  al. 2016) is ideal for obtaining good predictions by training the model using a conventional gradient descent algorithm with a sufficiently large training data set. we found that optimal results are obtained when training on at least 100  000 independent realisations, with performance severely decreasing when training on less than 10 000 realisations. the presented approach also has the major advantage of having limitless training data available, although a reasonable amount of training data should be chosen for computational feasibility. the network can also be trained to a larger grid than the 31 × 31 grid used here, giving a lot of flexibility for the specific model for which inference is needed. the 31 × 31 grid posed some issues for estimating ranges over a certain length. this limitation to the method can be remedied by increasing the grid size, but at the cost of increasing computation time during training of the cnn. to determine the right size of the grid in a practical case, we suggest training two preliminary networks with a small grid and a slightly larger grid. if predictions with these two networks deviate significantly when applied on real data, the grid size must be increased to accommodate all possible ranges. the process can be repeated iteratively until the same range interval is predicted for both networks, indicating a reasonable minimum grid size. using this strategy for the dk-model, a 31 × 31 grid was deemed suitable as showcased. our results in the synthetic case suggest that the deployed cnn has the same accuracy as a traditional automatic semivariogram fitting, but with a substantial improvement in speed, which now makes it feasible to analyse large grids within a reasonable amount of fig. 4 method results of a layer covering the danish island of fyn. (a) predicted ranges. (b) ranges clustered into regions. (c) layer elevation across the black profile shown in (a) and (b) and the predicted ranges (green/white line) across the same profile. the colour scale in (c) matches the colour scale in (b), showing where individual clusters are located. (a) range map 5.6 5.8 6 6.2 utm x [m] 6.08 6.1 6.12 6.14 6.16 u t m y [m ] ×106 0 1000 2000 3000 range [m] (b) cluster map 5.6 5.8 6 6.2 utm x [m] 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 cluster index (c) profile view 5.65 5.7 5.75 5.8 5.85 5.9 5.95 6 6.05 6.1 utm x [m] 0 20 40 60 80 100 e le va tio n [m ] 500 1000 1500 2000 2500 r an ge [m ] ×105 ×105 ×105 https://doi.org/10.34194/geusb.v53.8353 http://www.geusbulletin.org falk & madsen 2023: geus bulletin 53. 8353. https://doi.org/10.34194/geusb.v53.8353 7 of 7 www.geusbul let in.org computation time as showcased in the final validation example to account for non-stationarity. the subsequent clustering also makes it possible to define regions with comparable statistics in the stratigraphical model. for further application of the estimated statistical models, one could infer the local statistics from the sill and range estimates within each cluster and use these for, for example, localised geostatistical estimation (kriging) or simulation of each cluster instead of using a stationary model as done in madsen et al. (2022) for a hydrostratigraphic model. acknowledgements the authors would like to acknowledge the geological survey of denmark and greenland (geus) for providing funding for writing the manuscript as the method development was carried out in a consultancy project without funding for scientific publication. additional information funding statement the method was developed during consultancy work for the danish epa, whilst the hours for turning the results into a scientific contribution and writing the manuscript were provided by the geological survey of greenland and denmark (geus). author contributions faf: methodology, software, investigation, writing – original draft preparation rbm: conceptualisation, methodology, writing – reviewing and editing competing interests the authors declare no competing interests. additional files none provided references boisvert, j.b. & deutsch, c.v. 2011: programs for kriging and sequential gaussian simulation with locally varying anisotropy using non-euclidean distances. computers & geosciences 37(4), 495–510. https://doi. org/10.1016/j.cageo.2010.03.021 boisvert, j.b., manchuk, j. & deutsch, c.v. 2009: kriging in the presence of locally varying anisotropy using non-euclidean distances. mathematical geosciences 41, 585–601. https://doi.org/10.1007/s11004-009-9229-1 bongajum, e.l., boisvert, j. & sacchi, m.d. 2013: bayesian linearized seismic inversion with locally varying spatial anisotropy. journal of applied geophysics 88, 31–41. https://doi.org/10.1016/j.jappgeo.2012.10.001 cressie, n.a.c. 1993: statistics for spatial data. new york: wiley. hansen et al. 2013: sippi: a matlab toolbox for sampling the solution to inverse problems with complex prior information: part 1 – methodology. computational geoscience 52, 470–480. https://doi. org/10.1016/j.cageo.2012.09.004 higdon d., swall, j., & kern, j. 1999: non-stationary spatial modeling. in: bernado, j.m. et al. 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(eds.) advances in neural information processing systems 30, 7–8. new york: curran associates inc. https://proceedings.neurips.cc/ paper_files/paper/2017/file/32cbf687880eb1674a07bf717761dd3a-paper.pdf (accessed october 2023) https://doi.org/10.34194/geusb.v53.8353 http://www.geusbulletin.org https://doi.org/10.1016/j.cageo.2010.03.021 https://doi.org/10.1016/j.cageo.2010.03.021 https://doi.org/10.1007/s11004-009-9229-1 https://doi.org/10.1016/j.jappgeo.2012.10.001 https://doi.org/10.1016/j.cageo.2012.09.004 https://doi.org/10.1016/j.cageo.2012.09.004 https://doi.org/10.1016/0893-6080(91)90009-t https://doi.org/10.1016/0893-6080(91)90009-t https://doi.org/10.48550/arxiv.1602.07360 https://doi.org/10.48550/arxiv.1602.07360 https://doi.org/10.1007/s11004-021-09962-w https://doi.org/10.1016/b978-0-444-89178-5.50003-8 https://doi.org/10.1371/journal.pone.0266942 https://doi.org/10.1371/journal.pone.0266942 https://doi.org/10.1111/1365-2478.12848 https://doi.org/10.1016/j.enggeo.2022.106833 https://doi.org/10.1016/j.enggeo.2022.106833 https://doi.org/10.1068/a270615 https://doi.org/10.1068/a270615 https://doi.org/10.2113/gsecongeo.58.8.1246 https://doi.org/10.1007/s11004-023-10051-3 https://doi.org/10.22008/gpub/32631 https://proceedings.neurips.cc/paper_files/paper/2017/file/32cbf687880eb1674a07bf717761dd3a-paper.pdf https://proceedings.neurips.cc/paper_files/paper/2017/file/32cbf687880eb1674a07bf717761dd3a-paper.pdf https://proceedings.neurips.cc/paper_files/paper/2017/file/32cbf687880eb1674a07bf717761dd3a-paper.pdf machine learning-based estimation and clustering of statistics within stratigraphic models as exemplified in denmark introduction required resources methodological protocols producing training, validation and test data training the network validation validation on synthetic data validation by application of the method discussion and outlook acknowledgements additional information funding statement author contributions competing interests additional files references figures fig. 1 synthetic training data produced using gaussian semivariance models with random sill and range, with a component of noise. the circles show 120 randomly drawn points from each realisation. shading: smallest values are shown in blue and largest values are shown in yellow, with in-between values shown in green. fig. 1 synthetic training data produced using gaussian semivariance models with random sill and range, with a component of noise. the circles show 120 randomly drawn points from each realisation. shading: smallest values are shown in blue and largest values are shown in yellow, with in-between values shown in green. fig. 3 synthetic examples: 8 realisations are shown – each of their own gaussian semivariogram model with some component of noise. the semivariance geological survey of denmark and greenland bulletin 24, 2011, 96 pp. 1 geological survey of denmark and greenland bulletin 24 • 2011 the east greenland rifted volcanic margin c. kent brooks geological survey of denmark and greenland danish ministry of climate, energy and building 22 geological survey of denmark and greenland bulletin 24 keywords east greenland, north atlantic, rifted volcanic margin, large igneous province, lip, palaeogene, basalt, syenite, nephelinite, carbonatite, uplift. cover sundown over the nunataks in the main basalts (skrænterne fm) to the south of scoresby sund. camped on the glacier, the 1965 oxford university east greenland expedition travelled and collected from this area on foot, manhauling equipment on the sledge to the left. the expedition results were published in fawcett et al. (1973). frontispiece: facing page mountains of horizontally layered basalt flows rising to about 2000 m on the south side of scoresby sund. typical trap topography as found throughout most of the kangerlussuaq–scoresby sund inland area. chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, department of geoscience, aarhus university; minik rosing, geological museum, university of copenhagen; finn surlyk, department of geography and geology, university of copenhagen scientific editor of this volume: adam a. garde editorial secretaries: jane holst and esben w. glendal referees: dennis k. bird (usa) and christian tegner (dk) illustrations: eva melskens with contributions from adam a. garde digital photographic work: benny schark graphic production: kristian a. rasmussen printers: rosendahls · schultz grafisk a/s, albertslund, denmark manuscript received: 1 march 2011 final version approved: 20 september 2011 printed: 22 december 2011 issn 1604-8156 isbn 978-87-7871-322-3 citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 24, 96 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk or at www.geus.dk/publications/bull © de nationale geologiske undersøgelser for danmark og grønland (geus), 2011 for the full text of the geus copyright clause, please refer to www.geus.dk/publications/bull 3 44 abstract . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . historical background . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . note on stratigraphic nomenclature and radiometric dating . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . components of the rifted volcanic margin . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . th e precambrian and caledonian crystalline basement . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . th e pre-basaltic sedimentary rocks (the kangerdlugssuaq group) . . . . . . . . . . . . . . . . . . . . . . . . . . . th e lower basalts . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . th e main basalts . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . extent . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . structure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . stratigraphy . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . composition and genesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . radiometric ages and timing of volcanism . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . zeolite zones . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . comments on the plume hypothesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . late basalts and sedimentary rocks (the prinsen af wales bjerge formation etc.) . . . . . . . . . . . miocene basalts . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . off shore basalts . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . basalts of northern east greenland . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . long-distance tephra falls . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . gabbroic intrusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . th e skaergaard intrusion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . other gabbroic intrusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . ultramafi c plugs etc. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . felsic intrusions and volcanic rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . th e kangerlussuaq – blosseville kyst area . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . th e mesters vig area . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . th e kialeeq (kialineq) – kap gustav holm area . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . syenite genesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . nephelinites, carbonatites etc. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . nephelinite genesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . dykes and sills . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . hydrothermal activity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . geomorphology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . structure of the east greenland margin and its setting within the north atlantic region . . . . . th e coastal fl exure and its dykes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . segmentation of the margin . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . vertical and horizontal distribution of igneous material in the crust . . . . . . . . . . . . . . . . plate-tectonic patterns . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . vertical movements and erosion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . time frame for magmatism and margin deformation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . mechanisms of north atlantic rifting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . summary and perspectives . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . concluding remarks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . references . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . appendix . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . contents 5 7 7 11 13 13 14 16 19 19 20 21 21 24 25 27 27 27 28 29 29 31 32 32 34 36 39 41 45 47 49 50 53 54 60 61 64 64 65 66 67 67 69 70 72 74 75 76 93 5 brooks, c.k. 2011: the east greenland rifted volcanic margin. geological survey of denmark and greenland bulletin 24, 96 pp. the palaeogene north atlantic igneous province is among the largest igneous provinces in the world and this review of the east greenland sector includes large amounts of information amassed since previous reviews around 1990. the main area of igneous rocks extends from kangerlussuaq (c. 67°n) to scoresby sund (c. 70°n), where basalts extend over c. 65 000 km2, with a second area from hold with hope (c. 73°n) to shannon (c. 75°n). in addition, the ocean drilling project penetrated basalt at five sites off south-east greenland. up to 7 km thickness of basaltic lavas have been stratigraphically and chemically described and their ages determined. a wide spectrum of intrusions are clustered around kangerlussuaq, kialeeq (c. 66°n) and mesters vig (c. 72°n). layered gabbros are numerous (e.g. the skaergaard and kap edvard holm intrusions), as are underand oversaturated syenites, besides small amounts of nephelinite-derived products, such as the gardiner complex (c. 69°n) with carbonatites and silicate rocks rich in melilite, perovskite etc. felsic extrusive rocks are sparse. a single, sanidine-bearing tuff found over an extensive area of the north atlantic is thought to be sourced from the gardiner complex. the province is famous for its coast-parallel dyke swarm, analogous to the sheeted dyke swarm of ophiolites, its associated coastal flexure, and many other dyke swarms, commonly related to central intrusive complexes as in iceland. the dyke swarms provide time markers, tracers of magmatic evolution and evidence of extensional events. a set of dykes with harzburgite nodules gives unique insight into the archaean subcontinental lithosphere. radiometric dating indicates extrusion of huge volumes of basalt over a short time interval, but the overall life of the province was prolonged, beginning with basaltic magmas at c. 60 ma and continuing to the quartz porphyry stock at malmbjerg (c. 72°n) at c. 26 ma. indeed, activity was renewed in the miocene with the emplacement of small volumes of basalts of the vindtoppen formation to the south of scoresby sund. although the basalts were extruded close to sea level, this part of east greenland is a plateau raised to c. 2 km, but the timing of uplift is controversial. superimposed on the plateau is a major dome at kangerlussuaq. east greenland presents a rich interplay between magmatic and tectonic events reflecting the birth of the north atlantic ocean. it was active over a much longer period (36 ma) than other parts of the province (5 ma in the hebrides, northern ireland and the faroe islands) and contains a wider range of products, including carbonatites, and felsic rocks tend to be granitic rather than syenitic. as expected, there are many similarities with iceland, the present-day expression of activity in the province. differences are readily explained by higher production rates and the thicker lithospheric lid during the early stages of development in east greenland. the igneous and related activity clearly results from plate-tectonic factors, but the relationship is not understood in detail. in particular, the nature of the underlying mantle processes, primarily the presence or absence of a plume, is still not resolved. author’s address natural history museum of denmark, øster voldgade 5–7, dk-1350 copenhagen k, denmark. e-mail: kent2039@live.com abstract 66 fig. 1. the north atlantic igneous province, showing onshore and offshore basalts, dipping reflector sequences, the caledonian front and offshore magnetic anomalies. sites of the ocean drilling program (with prefix odp), its precursor the deep sea drilling project (dsdp) and geophysical traverses of the sigma project (t-i to t-iv) are also shown on the east greenland margin, h.c. & duncan 1996; holbrook et al. 2001). a small area of basalt in baffin island (canada) is not shown. sisz: south iceland seismic zone. tfz: tjörnes fracture zone. jmfz: jan mayen fracture zone. rey kja ne s r idg e c al e d o n ia n f ro n t t-ii t-i c al ed o ni an f ro nt ro ck all p lat ea u scandinavia ocean and deep sea drilling sites geophysical traversest-i onshore basalts offshore basalts seaward-dipping reflectors kangerlussuaq a24 a20 jmfz a24 svalbard tfz sisz odp 104 dsdp 38 dsdp 81 odp 918 odp 919 odp 409 odp 408 odp 407 disko ubekendt ejland ittoqqottoormiit/ scoresbysund tasiilaq kialeeq mesters vig umivik nuuk svartenhuk m oh ns r idg e knip o vich ridge lena trough g ak ke l r id ge 60°n 70°n 50°n 0°30°w 15°w 80 °n greenland uk shannon north sea irm ing er b as in c23 rc24 r c21 r c19 r c6r ? lofoten basin møre platform 500 km t-iv charlie-gibbs fracture zone bar en ts se a hold with hope t-iii faroe islands vøring plateau arctic ocean atlantic ocean odp 914–917, 989–990 7 introduction historical background in 1874, j.w. judd, a leading british geologist, wrote: “in the hebrides i will now show that we have supplied to us that great geological desideratum – a number of volcanoes so dissected by the scalpel of denudation as to constitute, as it were, a series of anatomical preparations, from which we may learn directly the internal structure of the piles, and obtain bases for reasoning on the causes to which that structure owes its origin.” (judd 1874, p. 232). how much more ecstatic he would have been had he been able to work on the equivalent rocks on the western side of the atlantic ocean, in east greenland, where exposures are orders of magnitude more extensive, more varied and more deeply dissected with vertical topography about three times greater due to subsequent greater uplift. moreover, outcrops are often scoured clean by recent glaciers, and obscuring vegetation and superficial deposits are frequently lacking. the uplift and consequent vertical exposure make the area ideal for an analysis of the interplay of magmatism and tectonism during continental extension and subsequent lithospheric rupture. this paper presents an overview of the east greenland part of the north atlantic igneous province, one of the earth’s largest igneous manifestations (fig. 1). the justification for this overview is that vast amounts of work have been published since previous reviews by wager (1934, 1947), deer (1976), noe-nygaard (1976), brooks & nielsen (1982a). in addition, nielsen (1987) dealt specifically with the alkaline rocks, and upton (1988) and saunders et al. (1997) covered the entire north atlantic igneous province. this work also points to the now voluminous literature of the east greenland part. at the present time, abundant research is continuing on the skaergaard intrusion, although details of this largely fall outside the subject of this paper. the north atlantic igneous province is best known for the extensive literature on its british part. the east greenland sector is much less known, although it is by far the most voluminous and diverse part of the province and deserves to be better documented. this review focuses on published work mainly pertaining to the area between kangerlussuaq and scoresby sund (kangertittivaq) which is clearly the most productive part of the north atlantic igneous province in terms of volumes of magma, with innumerable major intrusions and over 65 000 km2 of lavas attaining a maximum thickness of over 7 km. also included are the areas to the north of scoresby sund and to the south as far as tuttilik, as shown in fig. 2. it does not consider the offshore geology in detail, except where relevant to the overall picture. offshore geology can be found in larsen, h.c. (1984, 1990) and the two volumes from the ocean drilling program dealing with the drilling off south-east greenland (proceedings of the ocean drilling program, scientific results, vols 152 and 163). previously, east greenland was effectively out of reach of ordinary geologists due to its remoteness, its lack of any infrastructure and the almost impenetrable sea ice constantly drifting down from the arctic ocean. expeditions had to be major, self-sufficient undertakings which were often isolated for months or even years at a time. for example, l.r.wager led a year-long expedition here as late as 1935–1936, during which there was no contact with the outside world (wager 1937). today the main obstacles for scientists are financial. access in recent years has been by air, using stol aircraft on gravel airstrips as well as helicopters, thus eliminating the uncertainty of ship navigation in ice-filled waters. consequently, many individual geologists and small groups have been fortunate to work in the area (albeit generally under the umbrella of well-organised expeditions), and great strides have been made in the description and understanding of the palaeogene ge88 ology. moreover, mining companies have been active since 1986 (in fact since the 1950s north of scoresby sund), greatly expanding the logistical opportunities for scientists. the earliest studies of cenozoic igneous rocks in east greenland were those of nordenskjöld (1909), backlund & malmquist (1932) and krokström (1944). these publications dealt with rocks north of scoresby sund (see fig. 2) as there was reportedly an agreement that the land to the south was to be left to british geologists, who were establishing a presence there during the 1930s. these publications are now of historical interest only, but note that holmes (1918) obtained information for his classical description of basaltic rocks of the arctic region from nordenskjöld (1909). knowledge of the area has greatly expanded in recent years, largely due to major expeditions in 1994, 1995 and 2000 under the auspices of the danish lithosphere centre (see below). as many as 60 geologists in a season took part in these expeditions, working from remote field camps, usually in two-man groups. this review is not intended to dig deeply into the petrogenesis of the rocks – most of the references contain speculations as to the origins of the features described. probably, it is true to say that rarely has a consensus emerged and it is not the intention to overburden the literature with more speculation: conclusions tend to vary with time but observations are hopefully largely fixed. this paper serves as a guide to the currently rather voluminous literature and aims to answer the questions: ‘what’s there, where is it and when did it happen?’ only to a lesser extent is asked the question ‘why?’ although broadly connected with the genesis of the north atlantic ocean (e.g. brooks 1973), the east greenland rifted volcanic margin formed over a much more protracted period (c. 26 ma) in the palaeogene and produced magmas spanning an extremely broad spectrum of rock types such as tholeiitic and alkali basalts, nephelineand quartz-syenites, nephelinites and carbonatites, all contributing to making it an attractive research object. the rocks potentially contain a wealth of information on the interplay of processes of magma genesis, rifting, continental break-up and early ocean-floor formation; information extending in both time and space. although it was widely accepted for many years that the north atlantic igneous province is the surface manifestation of a deep mantle plume still active today in iceland (morgan 1971; brooks 1973; white & mckenzie 1989) and possibly the same which gave rise to the siberian traps (smirnov & tarduno 2010), the subject has recently become much more controversial (e.g. depaolo & manga 2003; foulger & natland 2003; foulger et al. 2005 reviewed by class 2008; foulger 2010). it is worth recalling that the general mantle plume hypothesis was so widely accepted at its time of proposal because it elegantly explained many observed and hitherto enigmatic features, e.g. hot spots, the chain of hawaiian islands, aseismic ridges and many more. more recently, seismologists have claimed to be able to identify the hot roots of both iceland and hawaii extending deep into the mantle (wolfe et al. 1997, 2009; ritsema et al. 2011). thus any rejection of the mantle plume hypothesis must provide very convincing alternative explanations. ultimately, it may emerge that both views have merit, in that several plumes have been shown to have a gabbroic signature (e.g. sobolev et al. 2000; stroncik & devey 2011), suggesting the presence in the source regions of subducted ocean floor as required by the plume opponents. whatever the truth, the enormous volumes of igneous rock produced over a very short period count as one of the major geological events in our planet’s history, possibly with global implications to the biosphere (e.g. eldholm & thomas 1993). an understanding of the causes and effects of such large igneous provinces is essential to our understanding of the earth system, and the presence or absence of mantle plumes remains a central question. the north atlantic igneous province, as it is generally known today1, is one of the largest igneous provinces 1 originally studied by holmes (1918), who called it the ‘brito-arctic province’, it has also been called the ‘thulean province’, a name to be discouraged as it has precedence for extensive proterozoic rocks in northern west greenland (wenk 1961). facing page: fig. 2. geology of the east greenland margin showing the crystalline basement (precambrian and caledonian), palaeozoic/ mesozoic basins, the extent of plateau basalts and the locations of the main intrusive centres. (of the following geological localities: aputitêq – aputiteeq; igtutarajik – ittutarajik; kialineq – kialeeq; nûgâlik – nuugaalik; pátûlâjivit – pattuulaajivit, the second name is given in modern spelling. see also fig. 10). 9 wiedemann fjord scoresby sund 72°n 16°w 69°n 66°n 75°n 24°w 24°w kangerlussuaq tuttilik tugtilik inland ice fig. 12 fig.9a fig.10 cenozoic basalt paleozoic/cenozoic sedimentary rocks precambrian/caledonian basement palaeogene intrusion n un at ak r eg io n blosse vil le ky st major fault settlement name of intrusion 100 km christian iv g letscher prinsen af wales bjerge urbjerget gunnbjørn fjeld nansen fjord ryberg fjord føhnfjo rd øfjord iceland kong oscar fjord nordvestfjord jameson land rødefjord r ø d ef jo rd gåse fjo rd lindsay nunatak trekantnunatakker batbjerg liverp o o l lan d ittoqqortoormiit kap brewster kap dalton danmark s tr æ de j.a.d. jensen fjord miki fjord jacobsen fjord fig. 14a fig. 8 fig. 14b 24°w 16°w 75°n 72°n 69°n 66°n 24°w32°w sorgenfri g letscher fig.9b kejser franz joseph fjord hold with hope wollaston forland sabine ø shannon myggbukta kap broer ruys kruuse fjord kialeeq kialineq noe-nygaard imilik kap gustav holm sulussuut sulugssut skaergaard skaergaard lilloise borgtinderne werner bjerge kærven kap simpson kap parry kangerdlugssuaq gardiner kærven nûgâlik (ersingerseq) ndr. aputitêq, pátûlâjivit igtutârajik kap edvard holm 1010 in the world (fig. 1; for an overview see upton 1988 and saunders et al. 1997). it has produced c. 10 × 106 km3 of magma (eldholm & grue 1994; holbrook et al. 2001). the greenland sector contains by far the largest amount of onshore, exposed magmatic material in the form of an extensive lava plateau, intense dyke swarms and numerous plutonic complexes. while the vast bulk of the lavas and dykes are tholeiites, and some of the intrusions have a similar composition, many major felsic intrusions also occur which are both underand over-saturated with silica, but mainly just saturated (i.e. syenites). in addition, subvolcanic complexes, dykes and extrusives belonging to the nephelinite/carbonatite suite are also found, possibly related to fractures or accommodation zones at a high angle to the continental margin (karson & brooks 1999). extensive offshore areas of igneous rocks also occur, both as extensions of onshore lava plateaux and as seismic seaward-dipping reflectors (larsen, h.c. & jakobsdóttir 1988), interpreted and subsequently confirmed by drilling as lava flows. the occurrence of these various rock types is clearly related not only to the horizontal movement of continental break-up, but is also intimately connected to vertical movements of the lithosphere, which are responsible for the pronounced present-day topography. thus the highest mountain within the arctic circle, the c. 3700 m gunnbjørn fjeld (fig. 2), is found here. l.r. wager, who had previously been close to the summit of mt. everest, led the anglo-danish first ascent of gunnbjørn fjeld and used the opportunity for important geological observations on the plateau basalts and underlying sedimentary beds (see courtauld 1936). uplift is thought to have continued into the neogene and is subject to an ongoing study (miss green project, geological survey of denmark and greenland), but its cause is still obscure. vertical movements, especially uplift, are difficult to quantify both in magnitude and timing, but have been addressed in recent years both by landscape analysis and fission-track dating. as noted above, this uplift of east greenland gives substantial vertical exposure providing great advantages to geologists. similarly, the exposure is superb. essentially, only ice fields and glaciers block exposure and where the rocks are seen they often present clean, glacially polished surfaces. this contrasts to many similar areas around the world that have subsided and become covered with younger sediments or thick regolith. the conjugate, eastern margin of the north atlantic igneous province (fig. 1), which includes the vøring plateau, faroe islands, rockall and hatton bank and the british hebridean province2, is not uplifted to the same extent as in east greenland; in fact it is largely submerged except for the faroe islands. consequently, it is not dissected to the same extent and only exposes a relatively limited range of rock types. furthermore, it was only active for a fraction of the period exhibited in east greenland. only the isolation of the east greenland sector has worked against the very numerous and comprehensive studies seen in the hebrides, which, with the notorious exception of the skaergaard intrusion, makes them the most studied igneous rocks in existence. previously, most geologists’ knowledge of east greenland would have been limited to the work of wager & deer (1939) and wager & brown (1968) on the skaergaard intrusion, which rapidly became the best documented igneous body in the world (see young 2003, pp. 309–326). lately, also the regional geology has come to the attention of geologists worldwide, largely due to the work of the now defunct danish lithosphere centre (larsen, h.c. et al. 1995; brooks et al. 1996; nielsen et al. 2001). moreover, in the 1980s and 1990s there was a worldwide surge of interest in large igneous provinces (lips) and what such provinces might tell us about the interior workings of the earth (e.g. the validity of the plume hypothesis and the nature of mantle convection), their role in planetary geochemical cycles and their possible role in global extinctions. the skaergaard intrusion became so well known not only due to the work of l.r. wager and w.a. deer, but also because of the fortunate juxtaposition of good exposure and uplift, creating a good stratigraphic section, and the fact that the intrusion crystallised as a closed system, i.e. with a single injection of magma and no extrusion of lava. only a broad outline of its history and research is given here, as any comprehensive description would require its own voluminous memoir. indeed, about 500 papers have been published on the intrusion, and a collection of abstracts can be found at http://www.skaergaard.org. although l.r. wager and 2 note that the hebridean province is far removed from the shelf edge: the true conjugate margins for southern/central east greenland are the faroe islands and the hatton bank. 11 his companions concentrated on the skaergaard intrusion, they also made remarkably fine maps and descriptions of the geology of the surrounding area (wager 1934, 1947), which they covered on protracted dogsledge trips (wager 1937). wager also provided the first descriptions of the geology from south of ammassalik (now tasiilaq) to scoresbysund (now ittoqqortoormiit3), a distance of over 800 km (fig. 2); descriptions and maps which have subsequently been shown to be essentially correct. in the 1970s and early 1980s, this work was extended by small university groups (e.g. brooks & nielsen 1982a). in the 1980s and early 1990s, collaboration with mineral exploration activity focused on mafic intrusions: skaergaard, kruuse fjord and kap edvard holm (bird et al. 1991, 1995; arnason et al. 1997a, b; arnason & bird 2000), which in turn provided a boost to the university studies. for an overview of the skaergaard mineral deposit, see nielsen et al. (2005). in the period 1994 to 2002, the danish lithosphere centre was active in the area with major, helicopter-supported expeditions in 1994, 1995 and 2000 (larsen, h.c. et al. 1995; brooks et al. 1996; nielsen et al. 2001). during these expeditions the prime targets were a comprehensive, flow-by-flow collection from eight composite sections of the lavas backed up by oblique stereoscopic aerial photography covering more than 1600 km of the mountain sides, a new collection covering the entire stratigraphy of the skaergaard intrusion and new collections from several other intrusions including the major kangerdlugssuaq alkaline complex (c. 800 km2). this field work was supplemented by major and minor element analysis, isotopic studies and an extensive 40ar/39ar dating programme, while the aerial photography allowed 3d photogrammetric interpretation of the architecture of the lava pile. the structure of the margin was investigated in an offshore, multi-channel, seismic survey (sigma project, holbrook et al. 2001) and several sites were successfully drilled by the ocean drilling program (leg 152: larsen, h.c. & saunders 1998; leg 163: larsen, h.c. & duncan 1996). this was a major data-gathering period, the results of which are likely still being digested, especially since the demise of the dlc and dispersal of interested researchers. likewise, mapping and collecting from the areas north of scoresby sund were made from the 1970s onwards under the auspices of the geological survey of greenland (presently the geological survey of denmark and greenland; see later sections). note on stratigraphic nomenclature and radiometric dating the stratigraphic nomenclature adopted in the following text and shown in fig. 3 is largely due to wager (1947), who divided the sequence into four main units. the kangerdlugssuaq group comprising the pre-basaltic sedimentary rocks (the sorgenfri and ryberg formations) was overlain by the lower basalts consisting of the vandfaldsdalen and mikis formations (mainly lavas) and the main tuffs (predominantly pyroclastics; now called the hængefjeldet formation). above this came the main basalts, comprising the bulk of the lavas along blosseville kyst and its hinterland, and finally the prinsen af wales lavas succeeding the main basalts. soper et al. (1976a) used a slightly different nomenclature. they called the entire volcanic succession the blosseville group and divided it into the same three formations of the lower basalts, plus the irminger formation, which corresponds to the main basalts of wager (1947) and this work. on the basis of very detailed work in the scoresby sund area and later to the south, these main basalts were subdivided by larsen, l.m. et al. (1989) into the magga dan, milne land, geikie plateau, rømer fjord and skrænterne formations (fig. 3) – note that the magga dan formation, which is of rather localised extent, has subsequently been incorporated into the milne land formation (l.m. larsen, personal communication 2002). the nansen fjord formation (currently thought to be the uppermost division of the lower basalts) was added later as a result of the danish lithosphere centre’s work in the nansen fjord area (larsen, l.m. et al. 1999). some uncertainty about this 3 as far as possible, spelling of place names follow the current orthography for the greenlandic language. geological names maintain the older spelling, as they were first reported. a complication arises because some place-name spellings are in the local east greenlandic dialect, others in the official west greenlandic. equivalent danish names may also occur. in the final instance, the names on the latest editions of the 1:250 000 scale topographic maps (kort & matrikelstyrelsen, denmark) are used. north of scoresby sund, higgins (2010) has comprehensively covered all place names. 1212 area remains, largely due to the gap caused by the 20 km wide christian iv gletscher (fig. 2). overlying the main basalts are the prinsen af wales lavas inland and the limited remnants of the igtertivâ formation on the coast, where the post-basaltic sedimentary rocks of the kap dalton and kap brewster formations are found. later, also the miocene vindtoppen formation of only restricted extent was discovered (story et al. 2004). the latter four formations are not shown in fig. 3. geology is a historical science, and an accurate knowledge of timing is essential. in the present context, correlations between igneous activity in different places and between this activity and plate-tectonic events require accurate radiometric dates. most of the older geochronological work on the east greenland rifted margin was done using the k-ar method, but this has subsequently proved to be unreliable. the history of dating of the kræmer ø syenite is taken as an example. geological investigations show that this is a single infig. 3. correlation of the lava formations in east greenland and on the faroe islands (from larsen, l.m. et al. 1999), showing stratigraphic divisions and ages from storey et al. (2007b). chemical types indicated by different colours, as labelled. stratigraphy of the faroe islands from passey & jolley (2009), but correlation in detail uncertain. position of the skaergaard intrusion in this stratigraphy from larsen, r.b. & tegner (2006). faroe islands east greenland, 68°n 0 1 2 3 4 km lo w er b as al ts lo w er s er ie s m id d le se ri es u p p e r s e ri e s volcaniclastic rocks 5 pr insen af wales formation skrænterne formation rømer fjord formation geikie plateau formation nansen fjord formation milne land formation mikis and vandfaldsdalen formations ryberg formation p re se n t a tl an ti c o ce an enni formation malinstindur formation beinisvørð formation unconformity p re -b re ak -u p s u cc es si o n 59.9 ± 0.7 ma 60.1 ± 0.6 ma 56.1 ± 0.5 ma 57.7 ± 0.5 ma 56.8 ± 2.3 ma 56.8 ± 0.7 ma 55.2 ± 0.7 ma 54.9 ± 0.7 ma hængefjeldet formation 55.1 ± 0.5 ma 55.4 ± 0.9 ma 55.1 ± 0.05 ma 55.1 ± 0.05 ma 55.0 ± 0.4 ma 54.9 ± 0.5 ma skaergaard intrusion 55.59 ± 0.13 ma m ai n b as al ts lopra formation 62–60 ma (urbjerg formation) cretaceous/ paleocene sedimentary rocks (lahars, tuffs, agglomerates) sy n -b re ak -u p s u cc es si o n morb-like flows alkaline lavas high-ti tholeiite sanidine-bearing tuff with age determination tholeiitic basalt picrite volcaniclastic rocks age determination cretaceous–paleocene sedimentary rocks 13 trusion, so a single crystallisation age is expected. noble et al. (1988) reported the results shown in table 1, which fell into two groups. it is clear that the authors have gone to great lengths to establish the true age, but we are left in a quandary: is the age c. 69 ma, c. 54 ma, or c. 49 ma as shown by the isochrons? (note: this locality is one of the few where data allow the calculation of isochron ages; in other cases, one has to make do with one or two k-ar ages which can hardly be assessed). tegner et al. (2008) obtained a 40ar/39ar date on biotite of 50.4 ± 0.1 ma, which is close to the two isochron ages and is also most reasonable geologically. the oldest age of noble et al. (1988) is unacceptable for geological reasons. this example shows that no confidence should be placed on k-ar ages although they are still frequently quoted in the literature about the east greenland rifted volcanic margin. a selection of newer and reliable age determinations from different parts of the region is presented in the appendix. the data come from several different sources and are mainly 40ar/39ar determinations. components of the rifted volcanic margin the precambrian and caledonian crystalline basement no detailed study of the precambrian basement in the kangerlussuaq area has been published. the area largely consists of quartzo-feldspathic orthogneisses and local paragneisses, in amphibolite, locally granulite, facies typical of such rocks in the north atlantic craton and other parts of greenland. these rocks are generally assumed to be the eastward continuation of the archaean craton on the northern side of the palaeoproterozoic nagssugtoqidian fold belt of west greenland (e.g. henriksen 2008, p. 30; nutman et al. 2008). a 207pb/206pb isochron on rocks from the skaergaard area published by leeman et al. (1976) gave an age of 2980 ± 20 ma (1σ). this age is consistent with k-ar ages by wager & hamilton (1964) and a rb-sr whole-rock isochron from kræmer ø giving a similar age (c.k. brooks, unpublished data), confirming that the basement in the kangerlussuaq area is archaean. in detail, these rocks are extremely complex as described by leeman et al. (1976), who distinguished at least three major events assumed to be earlier than their radiometric age. the events include deposition, deformation and metamorphism of interbedded volcanic and sedimentary units on a pre-existing basement that includes disrupted peridotite bodies, intrusion of granitic and trondhjemitic bodies, which in turn were strongly folded, intrusion of mafic as well as potassiumrich granites and pegmatites and finally a metamorphic event closing with pervasive shearing. the metapelites contain sillimanite, biotite, garnet and quartz, and crystallised at c. 650–700°c and 3–4 kbar pressure. some observations on the precambrian basement from the same area were given by bird et al. (1985). they noted deformed dolerite dykes and a prominent red shear zone which has acted as a major conduit for meteoric water at some unknown time, but prior to the intrusion of the palaeogene dykes. for details of the caledonian terranes of central and northern east greenland, the reader is referred to the group 1 8045 amphibole 60.2 ± 1.9 8041 biotite 64.2 ± 2.4 8026 amphibole 58.0 ± 1.0 8025 amphibole 63.1 ± 1.2 average (mineral ages) 60.4 isochron age 49.5 ± 0.06 group 2 8032 amphibole 54.9 ± 1.0 8029 amphibole 54.3 ± 1.0 8027 amphibole 54.4 ± 1.1 average (mineral ages) 54.5 isochron age 49.3 ± 0.2 sample number mineral dated result (ma) table 1. k/ar age data, kræmer ø† † data from noble et al. (1988). errors are 1σ. 1414 memoirs by higgins & kalsbeek (2004) and higgins et al. (2008), the paper by higgins et al. (2004), and the semi-popular works by henriksen (2008) and brooks (2009). to the south, the western boundary of the caledonian fold belt (fig. 2) is located somewhere in the unexposed area between scoresby sund and nansen fjord according to wager & hamilton (1964). however, a caledonian intrusive complex was located at batbjerg beyond the head of kangerlussuaq by brooks et al. (1976a, 1981). in many ways this rather enigmatic occurrence resembles the high-potassium intrusions of assynt in north-west scotland by containing pseudoleucite as well as unaltered leucite (brooks et al. 1975; gittins et al. 1980). the scottish intrusions likewise occur in the foreland region of the caledonian fold belt, being absent from the overlying metasedimentary thrust sheets. perhaps the question of the location of the caledonian front in this area needs to be re-assessed. a swarm of strongly alkaline dykes occurring in the western part of liverpool land (kalkdal) were originally suspected of being cenozoic in age, perhaps due to their freshness. they are rich in mica and have been called monchiquites. a thorough description of the dykes was made in an unpublished masters project (kofoed 1998), and they were shown to be of lower triassic age (ar-ar data: isochron age 245 ± 2 ma; plateau age 243 ± 5.17 ma). they are witness to an extensional event at this time. palaeogene igneous rocks north of scoresby sund largely intrude or overlie rift basins in the caledonian region. these basins are well described (e.g. surlyk 1990) and are filled with palaeozoic and mesozoic sediments, beginning with devonian deposits. to the west, basalts overlie an uneven crystalline surface of precambrian and caledonian age, as can be clearly seen in the inner parts of scoresby sund (e.g. føhnfjord and gåsefjord). magmas were thus exposed to contamination from quite different materials. it is clear that after the caledonian compression events, most of east greenland underwent protracted, but limited and probably intermittent, extension. in the kangerlussuaq area, basin formation apparently did not take place until the latest mesozoic to earliest cenozoic (although the oldest part of the succession is not exposed). the situation changed dramatically in the palaeogene when continental rupture occurred with the production of voluminous magmas and subsequent sea-floor spreading during c24r time which began at c. 57 ma and lasted 3 ma (westerhold et al. 2007). thus there was a repetition of the opening and spreading of the iapetus ocean (the precursor to the present-day atlantic ocean), which had formed the caledonian mobile belt 500 ma earlier, and this is a classical example of the wilson cycle (wilson 1966). the pre-basaltic sedimentary rocks (the kangerdlugssuaq group) both the preand post-basaltic sedimentary rocks are extremely important as they give a clear insight into the environmental conditions before and after the volcanic event, not least the vertical movements that would be expected to be associated with a mantle plume. the picture is complicated. the latest pre-basaltic sedimentary rocks do record an uplift, but this was short-lived and limited, as thick hyaloclastites are found in the base of the basaltic succession. the following account is largely confined to the basin located in the southern part of the province (fig. 2), just to the east of kangerlussuaq (miki fjord to nansen fjord), a basin which has been little studied until recently. three outliers are also included: at kap gustav holm, at kap brewster and the palaeogene sedimentary rocks underlying the basalts to the north (i.e. hold with hope, wollaston forland and sabine ø). the northern basalts lie on thick palaeozoic–mesozoic basins whose nature is well described elsewhere (e.g. surlyk 1990; stemmerik et al. 1993), and it is only the palaeogene sedimentary rocks that are briefly mentioned here. the post-basaltic sedimentary rocks at kap brewster and kap dalton are treated in a subsequent section on p. 28. studies of the sedimentary rocks (known as the kangerdlugssuaq group) and overlying volcanic sequence have been published by wager (1934, 1947), soper et al. (1976a, b), higgins & soper (1981), nielsen et al. (1981) and larsen, m. et al. (1996, 1999a, 1999b, 2001). traditionally, the sedimentary deposits under the volcanic rocks were divided into the sorgenfri formation, now known to be aptian to coniacian in age and first identified on the eastern side of the sorgenfri gletscher, and the ryberg formation, which is campanian to danian (fig. 4). the base of the succession is not exposed, so the oldest sedimentary rocks will be aptian or older. the earliest sedimentary deposits of the sorgenfri formation are estuarine sandstones. the basin then gradually deepened and became increasingly marine, first with marine sandstones, later abundant mudstones. deep-water conditions lasted 15 until the paleocene (danian) when uplift occurred, and marine sandstones reappeared in the upper ryberg formation. in the mid-paleocene, an extensive, up to 10–20 m thick quartz conglomerate sheet, the ‘schjelderup member’ (fig. 5a), was laid down in an extensive, braided river system over the entire area and is assigned to the overlying vandfaldsdalen formation. this was closely followed by basaltic volcanism. subsidence must then have been renewed, as thick, foresetbedded hyaloclastites indicate a water depth of >100 m. similarly, coarse polymict breccia units probably derive from submarine slides. the basin sediments attain a thickness of c. 1 km. the base of the sequence is not exposed in coastal regions where it is thickest, but elsewhere rests on archaean gneisses; the unconformity is thus diachronous. the top of the kangerdlugssuaq group is set at the base of the fluvial conglomerate known as the schelderup member. this is part of a return to siliciclastic sedimentation caused by uplift in the early paleocene and closely followed by the earliest volcanic deposits. these vertical movements were discussed by fig. 4. the pre-basaltic sedimentary succession, redrawn from larsen, m. et al. (2006), and the lower basalts, redrawn from peate, i.u. et al. (2003), showing the large proportions of fragmental material and the facies changes across the basin. correlation across the lower basalts is hampered by the lack of clear stratigraphic markers and the marked changes in facies. the age of the vandfaldsdalen formation relies on a correlation to lavas inland dated by the 40ar/39ar method (hansen et al. 2002). note the conflict with the age of the oldest basalts (as determined by radiometric dating and known basalt stratigraphy, see main text), implying that the transition from the ryberg formation to the vandfaldsdalen formation should be revised to the period 62–60 ma (see main text). ryberg fjord primary pyroclastics reworked pyroclastics hængefjeldet formation mikis formation facies change ~ 5 0 0 m kangerlussuaq basin lower basalts ~10 km nansen fjord j.a.d. jensen fjord ryberg formation sorgenfri formation m a in b a sa lt s lo w e r b a sa lt s milne land fm nansen fjord fm jacobsen fjord miki fjord vandfaldsdalen fm (62–60 ma) lava hyaloclastite and breccia pyroclastic rocks epiclastic rockshiatus marine sandstone marine mudstone fluvial sandstone (schelderup member) faroe unconformity dan maa cam san con tur cen alb apt se tha ypr p al ae o ge n e c re ta ce o u s 120 110 100 90 80 70 60 50 ma e ar ly la te e o ce n e p al eo ce n e 1616 clift et al. (1998) – but see the reservations of dam et al. (1999) – as manifestations of the arriving head of the supposed mantle plume. most of the exposed sedimentary rocks belong to the ryberg formation. rich fossiliferous horizons are reported (soper et al. 1976a; larsen, m. et al. 2006), including belemnites, shark’s teeth (hoch 1992), ammonites and plants (ginkgo and metasequoia have been identified). earlier workers were hampered by only having access to a few coastal localities, although l.r. wager examined inland localities during his mountaineering excursion. increased field coverage has recently revealed a much more complete sedimentary succession, albeit with several significant hiatuses. several more recent articles (larsen, m. et al. 1996, 1999a, b) are both extensive and up to date and essentially replace previous work. in a further study, larsen, m. & whitham (2005) suggested that the kangerlussuaq area was the source of thick sands, which are present in the faroe–shetland basin. this is a significant conclusion regarding petroleum prospectivity. while the onshore greenland rocks are overmature, traces of hydrocarbons suggest that other offshore basins with sub-basaltic sedimentary rocks are promising (jonk et al. 2005). a subsequent provenance study (larsen, m. et al. 2006) using whole-rock chemistry and dating of detrital zircons concluded that the kangerlussuaq and west shetland basins were fed from different sources. the pre-basaltic sedimentary rocks and those interbedded with the lowermost of the lower basalts were used by larsen, l.m. et al. (1999a) and vosgerau et al. (2010) as field analogues for other offshore basins around the north atlantic ocean, i.e. the vøring, møre, rockall etc. a problem remains regarding the timing of the transition from the pre-basaltic sedimentary succession (ryberg formation) to the oldest basalts (vandfaldsdalen formation). larsen, m. et al. (2006, fig. 2) showed this as lowest eocene, but this is the same age as the top of the main basalts as determined by radiometric dating and basalt stratigraphy (c. 55 ma, see below). this implies that the transition from the ryberg formation to the vandfaldsdalen formation should be revised to the period 62–60 ma. at kap gustav holm, wager (1934) reported c. 150 m of arkosic sandstones overlain by pillow lavas, hyaloclastite tuffs and epiclastic mafic tuffs (see also myers et al. 1993). the rocks are strongly hornfelsed, and fossil evidence only allowed wager to conclude that they were not older than cretaceous. myers et al. (1993, p. 260) reported that “late cretaceous and early tertiary lamellibranchs occur near the top of the sequence”. a tiny exposure of pre-basaltic sedimentary rocks occurs to the north at kap brewster. nøhr-hansen & piasecki (2002), the latest researchers to examine this exposure, concluded on the basis of dinoflagellate cysts that the lower part of the succession is late danian while the uppermost part is early selandian in age, corresponding to c. 58 ma. as the immediately overlying basalts belong to the milne land formation of the main basalts, there is a hiatus of several million years, although no angular unconformity is seen on seismic profiles (larsen, h.c. & marcussen 1992). the lower basalts (hansen et al. 2002) or their equivalents recorded in offshore boreholes (sinton & duncan 1998) have not been found in the scoresby sund area. on hold with hope, wollaston forland and sabine ø, small exposures of palaeogene sedimentary rocks have been described by nøhr-hansen et al. (2011). one of the interesting observations reported by these authors is that the exposures in haredal on wollaston forland may contain the paleocene–eocene thermal maximum which is well dated at c. 55.8 ma (charles et al. 2011), thus constraining all the overlying basalts to be younger than this. in fact, the topmost sedimentary rocks at this locality are biostratigraphically dated to “not younger than mid-ypresian” (nøhr-hansen et al. 2011, p. 63) which is compatible with the age of the oldest basalt (40ar/39ar age 55.02 ± 0.49 ma quoted by these authors from l.m. larsen’s personal communication 2008). the underlying sedimentary rocks are of early to middle campanian age, and the palaeogene sedimentary rocks show a hiatus that probably spans most of the selandian and thanetian. the presence of coarse sands and conglomerates free from basaltic clasts and abundant reworked cretaceous palynomorphs testify to an uplift immediately prior to basalt extrusion. the lower basalts as noted above, the entire basaltic sequence (i.e. lower and main basalts, fig. 3) has been called the blosseville group by soper et al. (1976a, b). the lower basalts (informal name) were first recognised by wager (1934, 1947) in the southern part of blosseville kyst from kangerlussuaq to nansen fjord. in this work, the lower basalts are taken to include both the lower basalts and the main tuffs of wager (1947) or the vandfalds17 dalen, mikis, jacobsen and hængefjeldet formations of soper et al. (1976a). the more recently introduced nansen fjord formation (larsen, l.m. et al. 1999) is also tentatively included as it has not yet had its affinities adequately documented, and the correlation across nansen fjord is still unclear. the main basalts as defined in this work (see below) are the irminger formation of soper et al. 1976a). it is clear that a formalised stratigraphy such as that established by larsen, l.m. et al. (1989) for the main basalts is also needed for the lower basalts, although rapid facies changes and lack of marker horizons complicate the picture. the lower basalts have been described by soper et al. (1976a, b) and nielsen et al. (1981), their chemical and isotopic relationships by holm (1988), fram & lesher (1997) and hansen & nielsen (1999), and their eruptive style by peate, i.u. et al. (2003). they have a maximum thickness of c. 2 km, with marked facies changes over their exposed area (fig. 4), which is not seen in the main basalts. furthermore, they contain some prominent picrite flows, which are rare in the main basalts. on the basis of very few samples, the lower basalts were first shown by brooks et al. (1976b) to differ from the overlying main basalts in both mode of occurrence and composition, and this has subsequently been confirmed. the lava flows are generally thin, compound pahoehoe flows (fig. 5b), often with well-preserved pahoehoe structures (fig. 5c), and hyaloclastites (fig. 5d). compound flows in the sense of walker (1972) are much more common than in the succeeding main basalts. see also passey & jolley (2009) for a description of these lava types in the faroe islands. breccias (probable debris flows), pyroclastic and epiclastic deposits and subaerial scoria cones are common. much of the observed explosive activity seems to be hydromagmatic in origin. in contrast, fragmental and near-vent facies are relatively rare in the overlying main basalts. in general, the lower basalts differ from the overlying main basalts in being much less voluminous or extensive, and in containing large amounts of pyroclastic and epiclastic material (35–50%). isotopic studies (holm 1988; hansen & nielsen 1999) also show that they are more contaminated with continental material. perhaps the later, more voluminous magmas overwhelmed the aqueous reservoirs and erupted through channelways armoured by the earlier passage of magmas, reducing their explosivity and contamination. the lower basalts also contain some prominent picrite flows and sills, especially in the lower part of the vandfaldsdalen formation. a photomicrograph of a picritic sill emplaced into lower basalts is shown in fig. 5e. furthermore, there are some interbedded sandstones in the lowermost parts of the lower basalts, one of which has been documented in detail by vosgerau et al. (2010). it has proved very difficult to obtain a reliable radiometric date for the lower basalts due to alteration by burial (prehnite-pumpellyite and greenschist facies: bird et al. 1986), but they are thought to be contemporaneous with the lowermost lavas drilled offshore southeast greenland (tegner & duncan 1999), with some onshore dykes to the far south (storey et al. 2007b) and with the lowermost lavas in the inland urbjerg area, which fall in the range 62–60 ma (formalised as the urbjerget formation: hansen et al. 2002). this would make them penecontemporaneous with the vaigat formation lavas of west greenland (storey et al. 1998) and with the lavas of the small isles of the hebrides (pearson et al. 1996) and place them in magnetochron c26r, well predating any identifiable ocean-floor spreading in the area. there appears to have been a hiatus of several million years between the eruption of these lavas and the overlying main basalts, although the transition is not well documented. a similar hiatus occurs in west greenland, offshore east greenland and the faroe islands, which allows a correlation of these localities (see fig. 3). in the hebrides, all volcanism appears to be earlier than 58 ma (saunders et al. 1997 and references therein; chambers et al. 2005) or 55 ma in the faroe islands. however, intrusive activity continued longer (e.g. to 51.5 ± 1.8 ma in the case of the granitic mourne mountains of ireland: meighan et al. 1988). extrusive activity also continued longer in west and east greenland, as described below. compositionally, the lower basalts are variable, with both primitive picrites (fig. 5e) studied from the point of view of melt inclusions in olivine by nielsen et al. (2006) and others which have become contaminated with continental crust (holm 1988; hansen & nielsen 1999), perhaps by the mechanism studied by blichert-toft et al. (1992), although their study was based on the later macrodyke. the most recent chemical and isotopic studies by hansen & nielsen (1999) identified iceland-like lavas contaminated with crustal material: amphibolite facies gneiss in the lower part and granulite facies gneiss in the upper part. the lavas are similar to those drilled offshore to the south on odp leg 152, where dacites occur (absent onshore). also, contamination is more dramatic and its sequence was the reverse (granulite facies contaminants initially, fol1818 c d b a person rucksack e 19 lowed upwards by amphibolite-facies contaminants). it is worthwhile to bear in mind that such contamination studies (including those referred to later) are hampered by the extreme heterogeneity of the likely contaminants (mainly precambrian rocks), a point stressed by jensen (1998). the compositional variation in the local gneisses has never been systematically studied, and authors seeking contaminants for basaltic magmas have used available, scattered data acquired for other purposes. it is thought (peate, i.u. et al. 2003) that the lower vandfaldsdalen and mikis formations formed shield volcanoes in the miki fjord area. these subsequently subsided and created space for epiclastic deposits eroding from new vent areas to the east in the j.a.d. jensen fjord area and also, although to a lesser extent, the nansen fjord area (figs 2, 4) where the deposits are largely pyroclastic. these two facies make up the hængefjeldet formation of previous authors. the formation of voluminous hyaloclastic, pyroclastic and epiclastic deposits (which include pillow lavas, breccias, pyroclastic surge deposits, bomb beds and tuffs) is due to three factors: (1) the presence of underlying water-saturated, high-porosity sediment (particularly coarse sands and conglomerates, as seen in both the sorgenfri and ryberg formations), (2) submarine activity and (3) the proximity to vents. for the lower two formations, these vents are thought to have been located to the north of miki fjord, and for the uppermost hængefjeldet formation in the j.a.d. jensen fjord area, where an extensive vent area has been described. thus according to peate, i.u. et al. (2003) the activity seems to have moved with time from the west at miki fjord to the east at nansen fjord (fig. 4). the transition to the overlying main basalts is poorly documented. it is largely concealed by the c. 20 km wide christian iv gletscher, a main outlet from the inland ice. it is in this area that the nansen fjord formation occurs, separated from the overlying milne land formation by a breccia horizon. an orthopyroxene-bearing suite promises to be a useful marker horizon (nielsen et al. 2001). similar lavas are found in the faroe islands (hald & waagstein 1983). the main basalts introduction this section deals with outcrops of basalts in the area between kangerlussuaq and scoresby sund (67–70°n). it covers the extent and structure of the lava pile and its stratigraphy, composition, genesis and timing of eruption. the basalts farther to the north in the hold with hope – shannon area are discussed in a separate section. there has been an immense augmentation of data on all the basalts since the oxford expedition (fawcett et al. 1973) began a second phase of work on them over 40 years ago, when they were only known from the pioneering work of l.r. wager in the 1930s. the basalt plateau has now been mapped, sampled systematically and in detail, analysed for a wide range of elements and isotopes and dated both biostratigraphically and radiometrically. previously, such basalt plateaux were seen as boring and uniform. with the benefit of an armoury of geochemical techniques and improved understanding of deep-earth systems, they are now seen as incomparable windows to the inner workings of the planet. the main basalts (so named by wager 1947, but as the irminger formation by soper et al. 1976a) are greater than 6 km in thickness, probably more than 7 km in places, and are exposed over a vertical distance of almost 4 km from sea level to the top of gunnbjørn fjeld (fig. 2). magnificent trap topography abounds, e.g. on the southern side of scoresby sund and inland nunataks (see frontispiece and fig. 6a) and is particularly enhanced when snow lies on the weathered interfacing page fig. 5. lower basalts and their base. a: sandstones of the ryberg formation overlain by a conglomeratic sheet (the ‘schjelderup member’), 1 m thick, which is identified as the ‘break-up unconformity’ by larsen, h.c. & saunders (1998, fig. 14). above this is the first columnar basalt flow. the unconformity at the top of the ryberg formation is quite small, but appears exaggerated due to the cross-bedding in the underlying sandstones. b: typical thin, vesicular compound flows in the mikis formation (person and rucksack for scale). c: exhumed pahoehoe toes in the vandfaldsdalen formation. d: hyaloclastite breccia with pillows in sødalen, from a 300 m thick unit with foreset bedding. chilled margins and radial structure of the pillows are clearly seen. e: photomicrograph of a picrite from sill in the lower basalts (c. 0.5 cm across, crossed polarizers). such rocks also occur in the main basalts, but rarely as fresh as this sample. 2020 flow horizons. along the southern blosseville kyst, trap topography is less evident or absent due to dyke intrusion and incipient metamorphism in this deeper part of the succession which causes flows and dykes to weather as one unit (fig. 6b). these lavas are the most extensive in the north atlantic area and dominate the 400 km coastline of blosseville kyst, imparting its forbidding black aspect, where a succession of high, vertical headlands rise from the inhospitable waters of danmark stræde (fig. 2). as a result of its reputation for impenetrable drift ice, strong currents, thick fogs and high cliffs, this has become greenland’s most feared coastline. the danish lithosphere centre, whose aim was primarily to investigate the basalt plateau, operated using helicopters in 1994, 1995 and 2000 measuring and collecting on a flow-by-flow basis from a comprehensive series of profiles through the entire lava plateau, backed up with airborne stereoscopic imagery (larsen, h.c. et al. 1995; brooks et al. 1996; nielsen et al. 2001). the cambridge arctic shelf project (casp) has also operated in the area using snowmobiles. their aim was probably petroleum-related, but no publications are forthcoming. in recent years, the miss green project (geological survey of denmark and greenland) has focussed on the evolution of the topography. the main basalts are regarded as contemporaneous with continental break-up. they have an age which corresponds to the earliest offshore magnetic anomalies, they are all reversely magnetised (so far as is known), they contain some compositions typical of oceanic spreading zones, and the great increase in magma productivity is consistent with such a major event. this means that the lower basalts are pre-break-up. magmatic activity at this time extended from eastern canada to the hebrides, but according to the widely acclaimed views of lawver & müller (1994), the plume was located under the middle of the greenland ice sheet. similarly, when the main basalts erupted, the supposed plume was still under the ice sheet, and it was not until at 40–30 ma that it crossed the coastline east of kangerlussuaq. these events are not immediately reconcilable with the classic plume model with its broad mushroom-like head and narrow tail as expounded by e.g. campbell & griffiths (1990), and special conditions must be invoked. moreover, other views on the likely position of the plume head exist. brooks (1973) placed it under kangerlussuaq at the time of basalt extrusion, for the simple reason that this was where the magmatic activity was highest, presumably the same reason for white & mckenzie (1989) to place it in the same location. in addition, a completely different plume track was proposed by forsyth et al. (1986), based on outcrops of volcanic rocks. it must be concluded that the location of the centre of the supposed plume 60–50 ma ago is in doubt, although its present position is widely agreed to be under vatnajökull in iceland. extent the basalt plateau is described by wager (1934, 1947), fawcett et al. (1973), brooks et al. (1976b) and larsen, l.m. et al. (1989) – the last is by far the most thorough description – and extends from scoresby sund to kangerlussuaq, covering c. 65 000 km2. it attains more than 6 km in thickness, perhaps more than 7 km locally (brooks & nielsen 1982a; pedersen et al. 1997), although it is thinner in the scoresby sund region (c. 3 km) and thins inland. however, its original extent was much greater as shown by minor basalt outcrops as far south as the kialeeq area (formerly kialineq in the geological literature) and kap gustav holm c. 200 km south of kangerlussuaq. fission track studies show that its thickness south of kangerlussuaq cannot have exceeded 3 km (hansen 2000). its original extent to the north is unknown, although a substantial basalt plateau exists in the hold with hope and wollaston forland areas (fig. 2 and discussed below). whether these dispersed areas were once united is unknown. sills and dykes in the intervening area suggest that surface flows may originally have been present here, too (e.g. larsen, h.c. & marcussen 1992; hald & tegner 2000). there is no doubt that igneous activity was present over the entire area, covering more than ten degrees of latitude. it extends offshore forming the shelf in some areas, and in others seaward-dipping reflectors occur directly offshore as a continuation of the coastal flexure (fig. 1). they eventually pass into iceland-type oceanic crust (larsen, h.c. & saunders 1998). it is considered unlikely that the basalts pass under the inland ice and connect with those around disko bugt in west greenland, as the basalts are everywhere seen to thin inland, as indeed do those on the west coast. nevertheless, according to dawes (2009) there is still a probability that basalts underlie the inland ice, though the rationale for this is unclear. 21 structure the general impression of the basalt area is one of approximately flat-lying flows forming the classic trap topography (see frontispiece and fig. 6a). many of these flows are of immense size: up to hundreds of cubic kilometres in volume (larsen, l.m. et al. 1989; pedersen et al. 1997). in detail, however, the plateau has been deformed as described by pedersen et al. (1997), who used multi-model photogrammetrical analysis using the images obtained in 1995. these authors identified the following features, in part confirming previous reports: (1) an inland plateau with lava dips of <2° (fig. 6a). (2) a coastal flexure zone, associated with an intense dyke swarm, with dips of 30–60° towards the ocean (as originally described by wager & deer (1938), and seen in fig. 6c). (3) a prominent, heavily fractured zone, c. 10 km wide and trending roughly n–s, approximately colinear with a major caledonian fault in the area north of scoresby sund. this may represent an abortive location of break-up, used by the earlier basalt formations before an eastwards jump to the present margin (larsen, l.m. & watt 1985). (4) a local point source uplift c. 50 km inland where lavas are fragmented into blocks dipping centripetally. this structure, although larger, is perhaps similar to the one described by matthews (1979) in the watkins bjerge area, which he thought might represent the roof of an unexposed intrusive body. (5) a half-graben preserving mildly alkaline basalts (igtertivâ formation) above the main lava pile in the northern blosseville kyst area. (6) a structurally complex zone of tilted blocks paralleling the coast but inland of feature 2 and bordered by feature 5 to the north-west. (7) the kangerlussuaq dome as originally recognised by brooks (1979), now heavily dissected, but with eroded basaltic flanks still reaching up to c. 3700 m altitude. naturally, flows vary considerably in thickness and morphology, a fact helpful to mapping. thus distinct packages of flows, identifiable by characteristic thicknesses, colours etc., can sometimes be followed from nunatak to nunatak. in general, these flows are simple flows in the sense of walker (1972), implying very much higher extrusion rates than for the largely compound flows of the underlying lower basalts (figs 5b, 6b). nevertheless, compound flows do occur, as do limited hyaloclastites, especially in the north where flows have filled valleys in the underlying gneisses. some hyaloclastites also occur in the uppermost lavas (skrænterne formation) that may have formed in lakes. red boles (fossil soil horizons, fig. 6d) are not common except towards the top of the pile, reinforcing the observation from radiometric dating and biostratigraphy that the eruptions followed each other rapidly, only slowing towards the top of the succession. likewise, intrabasaltic sedimentary rocks have only been found close to the top of the succession. their marine fossils (fig. 6e) show that the thickening of the lava pile was balanced by deepening of the basin. common textures are aphyric to microphenocrystic, although many flows with large phenocrysts of plagioclase, often forming star-shaped aggregates, are quite common. vesicular flow bases and pipe amygdules are common (fig. 6f). picrites are relatively rare compared to the lower basalts and the overlying prinsen af wales lavas. stratigraphy the stratigraphy of the main basalts (fig. 3) was originally established in the scoresby sund area by larsen, l.m. et al. (1989) after many years of detailed field and laboratory work. mapping of the basalts requires sophisticated techniques as flows can rarely be followed far and marker horizons are few, with the exception of the distinctive hjørnedal marker horizon (larsen, l.m. et al. 1989). otherwise, correlations are established by distinctive flow packages and backed up by sampling on a flow-by-flow basis and chemical fingerprinting. certain lava types and successions of types may have quite distinctive chemical characteristics. this work established six formations, of which two, magga dan and igtertivâ, are not regionally developed. as already mentioned, more recent work to the south has identified a seventh formation: the nansen fj ord formation, which is c. 1300 m thick in the nansen fjord area, 2222 fig. 6. main basalts. a: the basalts inland (height of wall about 1100 m) at borgtinderne intruded by the borgtinderne syenite in the lower spurs (brown et al. 1978). here, the flows are almost horizontal, lacking the seaward dip (the coastal flexure) seen in coastal areas. facing page: b: cliffside in miki fjord (height c. 500 m) showing massive appearance of basalts due to incipient metamorphism. photo from eskimonæs by a.a. garde. c: cliff in rømer fjord typical of the main basalts, viewed towards north. the dip of the lava flows is due to the coastal flexure. both simple flows (thick, rusty-coloured units) and compound flows (thinly bedded, greyish brown units – at upper right) are present. dykes (arrows), approximately perpendicular to the flows and with bleached margins dip inland, also as a result of the flexure. height of cliffside c. 100 m. d: detail of a red bole (fossil soil horizon, possibly weathered tephra) and the overlying basalt flow at kap dalton, with chilled lower contact of the flow and lower part of colonnade (prismatically jointed) zone. the author for scale. e: specimen of the crab, coeloma bicarinatum, a distinctive eocene fossil from the post-basaltic kap dalton formation (sample 9 cm across). f: lower part of a basaltic flow unit showing vesicular base and pipe amygdales above. white material is zeolite. g: main basalts at wiedemann fjord, little affected by flexuring. they are cut by a large dyke filled with hydrothermally deposited minerals, mainly calcite with minor copper sulphides. note its continuation in the distant mountains (arrow). but which is now assigned to the lower basalts. it is divided from the overlying milne land formation by a prominent sedimentary horizon that itself rests on paleocene sedimentary rocks or gneisses inland. this horizon can be correlated to the faroe islands (larsen, l.m. et al. 1999 – see also søager & holm 2009). the relationship of the nansen fjord formation, from where it was established to the east of nansen fjord to the exposures on the western side of nansen fjord, is unclear, and thus there is a corresponding lack of clarity regarding the precise relationships of the succession to the underlying lower basalts on each side of the fjord. radiometric dating suggests a substantial hiatus of several million years but not yet clearly defined. according to larsen, l.m. & watt (1985), the main basalts comprise two major cycles. the lowest and variable milne land formation is succeeded by the relatively uniform geikie plateau formation. above this comes the variable rømer fjord formation and, above this, the uniform skrænterne formation. this division at the base of the rømer fjord formation was also apparent in the work of tegner et al. (1998a) where it was seen to signal a change in the melting regime. the main difference between the two cycles is that the lower one is largely quartz normative, while the upper is overwhelmingly olivine normative, suggesting reduced crustal influence with time. in the scoresby sund area and inland, the basalts rest on precambrian or caledonian gneisses where a considerable palaeo-topography of c. 300 m exists (seen in fig. 5 of larsen, l.m. et al. 1989). this relationship is perhaps best seen in gåsefjord and føhnfjord in the interior of the scoresby sund fjord system (fig. 2), where the underlying rocks are the krummedal paragneisses of the niggli spids and hagar bjerg caledonian nappes and the precambrian foreland (higgins et al. 2004). 23 g f e b c d 2424 composition and genesis the predominant composition of the lavas is described as high-tio2 ferrotholeiites. picrites also occur in minor amounts, often found in certain horizons, besides orthopyroxene-phyric ankaramites interpreted as products of magmas contaminated with continental crust (gras et al. 2004) and low-tio2 flows interpreted as similar to morb. in general, it seems that the lowermost part of the post-break-up succession is most variable. in the sydbræ area of scoresby sund, a nephelinite tuff, up to 12 m thick, was identified by larsen, l.m. et al. (1989) at the base of the rømer fjord formation. as it does not occur in other profiles, it is probably quite local; larsen, l.m. et al. (1989) stated that it covers 20 km2. this and a thin sanidine-bearing tuff in the uppermost skrænterne formation are unique as all other extrusives in the main basalts are tholeiitic, except for the uppermost, very restricted, igtertivâ formation which is poorly characterised but known to include some alkaline flows. the sanidine-bearing tuff at skrænterne is thought to be sourced in the gardiner complex (fig. 2) and is extremely extensive, and has proved useful for radiometric dating (heister et al. 2001 and see below). tegner et al. (1998a) showed that the main basalts can be divided into minor, low-titanium flows (0.82–1.96% tio2), those which resemble morb, and high-titanium flows (1.63–6.18% tio2) – a type that is widespread in the north atlantic igneous province. their key diagram is reproduced here as fig. 7. the former type, regarded as being morb-like, was thought to reflect the early phases of sea-floor spreading, while the latter group has been subdivided into a high-ti group and a very high-ti group. two major units make up the sequence (fig. 7). tegner et al. (1998a) discussed a genetic model in which the high-ti lavas of the milne land and geikie plateau formations were produced by segregation at decreasing temperatures whereas those of the rømer fjord and skrænterne formations segregated at near-constant temperature but over a wide range of pressures. alternatively, the low-ti lavas, which resemble mid-ocean ridge basalts, may simply have originated in local, depleted regimes of the mantle rather than being sourced at a true ocean ridge. in any case, their presence shows that the source region for the main basalts was heterogeneous. andreasen et al. (2004) made a flow-by-flow study of the relatively homogeneous geikie plateau formation, identifying fractional crystallisation as the dominant process of variation and proposing a model for mg-number sf rff gpf 0 30 40 50 60 0 2 4 0 1 26 0.8 1.2 1.6 2.0 100 200 300 mlf tio2 (wt.%) la/smn dy/ybn fl o w n u m b er in c o m p o si te la va p ro fil e lo w er p o rt io n u p p er p o rt io n fig. 7. chemical stratigraphy of the plateau basalts based on 331 flows covering a vertical section of 6 km. red symbols: lower unit of high-ti lavas. blue symbols: upper unit. yellow symbols: low-ti lavas. from tegner et al. (1998a); reprinted by permission from nature, macmillan publishers. 25 the plumbing system. contamination by amphibolitefacies gneiss in the main flows and granulite-facies gneiss in the evolved flows was minor, but indicated that the differentiation took place in relatively shallow magma chambers. they suggested that these lavas are contemporaneous with the skaergaard intrusion, although peate, d.w. et al. (2008), in an extension and refinement of this study, pointed out that pb, nd and sr isotopic data indicate that the skaergaard magma is more contaminated with crust. this correlation was later used in an estimate of the timing of the basalt extrusion relative to the skaergaard intrusion by larsen, r.b. & tegner (2006) as noted below. peate, d.w. & stecher (2003) studied a suite of least contaminated lavas representing the main types and determined their pb isotopic signatures. they found that the samples were indeed minimally contaminated with continental crust, and that there was a correlation with type (e.g. 206pb/204pb increasing from low-ti to very high-ti, mirrored by εnd). they concluded that the east greenland volcanics sampled source regions similar to those now being tapped in iceland. a comparison of the east greenland basalts with those of iceland was undertaken by barker et al. (2006) to investigate in more detail the nature of the mantle sources being tapped through time. many recent studies have shown that the icelandic sources are heterogeneous (e.g. thirlwall et al. 2004) and can be explained by mixing of several sources enriched and depleted in pb and nd space. barker et al. (2006) found that indeed the greenlandic lavas show similar components, although differing in detail, allowing the authors to make some inferences on the nature of the mantle plume, which they thought best explained the volcanicity. precious metals in east greenland basaltic rocks were first investigated by nielsen & brooks (1995), who found elevated levels in some rock types. later, two detailed studies (momme et al. 2002, 2006) catalogued the platinum-group element (pge) behaviour in a representative suite of main basalts. combining their pge results with concentrations of rare-earth elements, they were able to show that the low-ti lavas had formed by high degrees of melting (c. 20%) in a cylindrical melting regime of the mantle, while the high-ti lavas had formed from a more sulphur-poor source by lower degrees of melting (c. 6%) in a mantle melting regime of triangular shape. three additional suggestions for the origin of the distinctive features of the main basalts (i.e. enrichment with respect to morb basalts) should be noted. bernstein et al. (2001) suggested that the enriched component in these basalts may represent melts of recycled ocean crust, which could be represented by the inland occurrences of melilitites and nephelinites (see below). only c. 7% of the nephelinite needs to be incorporated into normal mid-ocean ridge basalt. the suggestion that recycled oceanic crust may be involved is, of course, favourable to the idea that formation of the basalts involved melting of a subducted slab, as contended by the opponents of the mantle plume hypothesis (foulger et al. 2005). gabbro signatures suggesting earlier subducted material have been identified at other postulated plume locations (e.g. sobolev et al. 2000; stroncik & devey 2011). studies by hogg et al. (1988, 1989) and larsen, l.m. et al. (1989) showed that substantial compositional variation within the basalts can be accounted for by refilling-tapping-fractionation processes as described by o’hara & matthews (1981). finally, in a study of the composition of likely primary melts, korenaga & kelemen (2000) showed that the north atlantic source regions have been heterogeneous in terms of fe/mg ratios since the inception of activity up to the present day. much compositional variation is thus attributable to variation in the source of the magmas. radiometric ages and timing of volcanism biostratigraphic methods relying on the pre-basaltic ryberg formation and the post-basaltic kap dalton formation established that the basalts were emplaced in the interval between the latest paleocene and the earliest eocene (soper et al. 1976a, b). at that time, radiometric dating was difficult and results inspired little confidence. worse, there was little chance of appraising the quality of the results. the only method available at that time was k-ar and this suffered from variable and indeterminate ar loss from altered material (probably mesostasis glass). in other cases, excess ar was present (i.e. some 40ar had been present from the start, due to imperfect degassing). the first k-ar results were reported by beckinsale et al. (1970) and later by noble et al. (1988), not only on basalts, but also on the intrusions (see p. 13). results spread so widely that little conclusion could be reached except that the basalts were paleocene/eocene, and this was already known from biostratigraphy. 2626 the 40ar/39ar method is both more accurate and allows an appraisal of the quality of the results using internal checks (e.g. nature of plateau age, correspondence between plateau and isochron ages, approach to atmospheric values for common ar). its main weakness is that it relies on calibration to known samples rather than yielding an absolute age as do u, th-pb methods. the first attempt to use the 40ar/39ar method was undertaken by hansen et al. (1995) with some, but limited, success. extensive 40ar/39ar studies (storey et al. 2007a, b) show that the age range of the main basalts is c. 58 to 55 ma; results are shown in fig. 3. the lower age is corroborated by a date of c. 55 ma, reported by heister et al. (2001), for the sanidine-bearing tuff near the top of the skrænterne formation (fig. 3), which is the same as that obtained by storey et al. (2007a) on the same material. the age data show that the basalts were extruded at around the time of continental break-up in this area, consistent with the presence of magnetic anomaly 24r offshore and the occurrence of morbtype lavas within the succession. as noted previously, chron 24r is centred at c. 56 ma with a duration of c. 3 ma (westerhold et al. 2007). the sequence from the lowest milne land formation to the uppermost skrænterne formation extends from 56.1 ± 0.5 to 55.0 ± 0.4 ma, confirming the suspicion that this >6 km thick sequence of lavas was erupted very rapidly, as has been found in other large igneous provinces (e.g. the deccan: chenet et al. 2007; yemen–ethiopia: coulie et al. 2003). unfortunately, the results do not rigorously define the length of the hiatus marked by the sedimentary/pyroclastic horizon at the base of the milne land formation, and which is present as coal and volcaniclastic rocks in the faroe islands. two samples from the underlying nansen fjord formation give 56.8 ± 2.3 and 57.7 ± 0.05 ma. combined with determinations from the urbjerget formation (hansen et al. 2002) and the faroese lower series (storey et al. 2007a, b), the age of the lower basalts in east greenland is c. 60 to c. 57 ma, and the earliest milne land formation flows are c. 56 ma. the best estimate for the length of the hiatus is only c. 1 ma, but this may yet be revised as no study has specifically focused on this problem. storey et al. (2007b) discussed the rate of magma production in the east greenland –faroe sector of the north atlantic igneous province. they showed that volcanism began at 61 ma and continued to just after 57 ma over an area 2000 km across from baffin island to britain. this early stage is represented by the lower basalts. an average production rate of 163 km3/ma per kilometre of rift length was estimated (allowing for estimated underplated and intruded material), although this decreased slightly with time. after the volcanic hiatus at the base of the milne land formation in east greenland and the base of the middle series in the faroes, this value jumped to 3000 km3/ma per kilometre of rift length, later decreasing to half of this. essentially, this is in agreement with the seismic results from offshore east greenland reported by holbrook et al. (2001), based on the traverses shown in fig. 1. storey et al. (2007b) considered models of edge-driven upwelling, meteorite impact and start-up mantle plume to explain these amazing rates of magma production and concluded that only the plume model is a likely explanation. in a novel argument, larsen, r.b. & tegner (2006) used estimated pressures in skaergaard gabbro pegmatites to show how the thickening of the overlying plateau basalts had taken place during the evolution of the skaergaard intrusion, confirming the very rapid rate of several kilometres of lava extrusion, perhaps taking as little as 300 000 years. note that red boles (or intra-flow weathered tephras), which indicate substantial periods between eruptions, are rare in the lower parts of the main basalts, although specific observations are undocumented. that shown in fig. 6d is near the top of the pile when extrusion rates were apparently waning. sporadic palaeomagnetic observations (tarling et al. 1988; abrahamsen & nordgerd 1994) also supported a rapid extrusion, as all the lavas examined so far are reversely magnetised and were therefore probably emplaced during magnetochron c24r. according to westerhold et al. (2007), who used astronomical tuning in addition to radiometric methods, magnetochron c24r lasted 3.118 ± 0.05 ma. note, however, that palaeomagnetic studies only cover a short section of the pile. there are now reliable radiometric ages for the main basalts and also for the sanidine-bearing tuff in the skrænterne formation (shown in fig. 3). heister et al. (2001) argued that this tuff is represented at many distant sites including deep sea drilling project site 550, where it occurs just above the paleocene–eocene thermal maximum, albeit with an intervening hiatus in deposition. passey & jolley (2009) suggested that the nansen fjord formation and its equivalent in the faroe islands post-date the paleocene–eocene thermal maximum. this could possibly be true, although the dating and rates of lava extrusion still lack the necessary precision and accuracy to support this. larsen, r.b. & tegner (2006) estimated that 5.3–6.3 ± 2.7 km of basalt stratigraphy were produced in less than 27 300 000 years, while westerhold et al. (2007) estimated that 672 000 years had elapsed between the inception of the paleocene –eocene thermal maximum and ash layer –17 in denmark, which is correlated to the sanidine-bearing tuff layer at the top of the skrænterne formation. thus the entire c. 7 km of lava could easily have been formed in this interval. however, storey et al. (2007a), relying on their radiometric dates, especially for the sanidine-bearing tuff originally dated by heister et al. (2001), place the paleocene–eocene thermal maximum in the geikie plateau formation and the associated carbon isotope excursion around the top of the milne land formation. on a separate, but related note, the claim by jolley et al. (2002) that the geological time scale for the palaeogene is seriously in error should be ignored for the reasons stated by aubrey et al. (2003), wei (2003) and others. the long-standing question as to the possible connection between the paleocene–eocene thermal maximum and volcanism in the north atlantic region (e.g. svensen et al. 2004, 2010; hansen 2006) is very significant in the light of the present interest in the climatic changes and will be addressed below. the best estimate range for the age of the paleocene–eocene thermal maximum is currently 55.964 to 55.728 ma (charles et al. 2011). zeolite zones zeolites were reported by fawcett et al. (1973) and larsen, l.m. et al. (1989), but the most comprehensive treatment was presented by neuhoff et al. (1997). they recognised the following zones with depth: no zeolites (0 to c. 400 m), chabazite + thompsonite (c. 400 to c. 1400 m), analcime (c. 1400 to c. 1600 m), mesolite + scolecite (c. 1600 to c. 2250 m), heulandite + stilbite (c. 2250 to c. 2500 m) and laumontite (down to c. 2650 m). below this, prehnite + k-feldspar + quartz are stable. they correlated these zones with similar ones in iceland and were able to set temperatures on the different isograds that are consistent with a geothermal gradient of c. 40 ± 5°/km. detailed mapping showed that zeolite isograds transgress the stratigraphy but are earlier than both plutons and the dyke swarm, and reflect synvolcanic deformation. the lack of zeolites in the uppermost 400 m suggests that the top of the lava pile inland is close to the original palaeosurface, i.e. erosion has been minimal here. comments on the plume hypothesis since the present author first proposed the application of the plume hypothesis to the area (brooks 1973), numerous attempts have been made to confirm this idea, but in spite of all the information now available, the question is still unanswered. most workers agree that there are close ties between the east greenland main basalts and modern icelandic lavas, in particular evidenced by the similar isotopic signatures. it is still the consensual opinion that the vast outpourings of lava over a very short period and the continuing activity over close to 60 ma are best explained as the result of a plume, where the initial huge outpourings represent a mushroom-like head, and continuing activity in iceland is sustained by the plume stem, following the classic model of campbell & griffiths (1990). it should be remembered that when the plume theory was first discussed in the early 1970s, the internal structure, composition and age distribution of flood basalt provinces were poorly known, such provinces not then being attractive research targets. the plume hypothesis changed this situation radically, and large igneous provinces rapidly became a major focus of earth scientists. nevertheless, such provinces remain enigmatic and are not explained by standard platetectonic scenarios. furthermore, the suggestion that plumes have been associated with global extinctions as championed by several workers (e.g. eldholm & thomas 1993; courtillot 2002) has gained increasing support and attracted researchers from many fields. east greenland, with its outstanding exposure of part of one of the earth’s largest igneous provinces, represents an unequalled opportunity for the study of this hypothesis. late basalts and sedimentary rocks (the prinsen af wales bjerge formation etc.) originally described by wager (1947) and anwar (1955) and more recently by hogg et al. (1988, 1989), brown et al. (1996, 2000), hansen et al. (2002) and peate, d.w. et al. (2003), a series of variably dipping lavas occur in the nunataks of the prinsen af wales bjerge (fig. 2). they sometimes show angular discordance with the underlying plateau basalts and are thought to be late volcanic edifices unconformably overlying the plateau 2828 basalts. lavas also occur on the trekantnunatakker situated in the middle of kangerlussuaq gletscher (fig. 2). however, anwar (1955) and peate, d.w. et al. (2003) showed that these lavas are partially intercalated with the uppermost plateau lavas. they suggested that there is no sharp transition between the main basalts and the overlying prinsen af wales formation, the latter originating from more inland, central eruption sites, which became active as the production of the plateau lavas (main basalts) was waning. these basalts are alkaline and often strongly olivine and pyroxene-phyric: very different from the underlying main basalts, which are tholeiitic and typically aphyric to plagioclase-phyric. some of these flows are basanites and melanephelinites (nielsen et al. 2001), representing the type of magma from which the gardiner complex evolved. several eruption sites have been identified accounting for the tilted flows, i.e. they are draping the sides of a volcanic edifice. on lindsey nunatak (west of the prinsen af wales bjerge, fig. 2), there are intervals of polymict conglomerates containing a variety of clasts, most notably some containing eudialyte, a mineral otherwise only known from the gardiner complex (nielsen 1989). also on this nunatak, the prinsen af wales bjerge lavas are overlain by picritic tholeiites whose affinity is unknown. there is some evidence that the sedimentary rocks here were deposited in rivers running west, perhaps down the side of the emerging dome (see geomorphology section). peate, d.w. et al. (2003) used sm-nd-pb-hf-os-he isotopic data to show that the alkaline lavas are more radiogenic than typical icelandic lavas and suggested they may contain recycled material (as has also been suggested by bernstein et al. (2001) for the enrichment in incompatible elements in the ti-rich lavas of the underlying plateau basalts). also overlying the plateau basalts in the northern coastal areas and of very limited extent is the igtertivâ formation (larsen, l.m. et al. 1989), known only as downfaulted remnants in a half-graben at kap dalton on blosseville kyst south of scoresby sund (fig. 2). additional outcrops may be awaiting discovery elsewhere, as the area is little visited. clearly, the igtertivâ formation was originally much more extensive. these lavas also include alkaline varieties. sedimentary rocks overlie the basalts unconformably at both kap dalton and kap brewster (fig. 2). they have been described by ravn (1904), wager (1935), birkenmajer (1972), larsen, l.m. et al. (1989) and larsen, m. et al. (2005). the lower of the two successions, formerly the kap dalton formation, is now known as the kap dalton group. this comprises the lower bopladsdalen formation (formerly the cyrena beds) and the upper krabbedalen formation (formerly the coeloma beds). these deposits provided a biostratigraphic upper limit of ypresian to the plateau basalts (soper & costa 1976). the basal conglomerate of the bopladsdalen formation contains clasts of a wide variety of strongly alkaline rocks of unknown provenance (wager 1935). the krabbedalen formation contains a rich, shallow marine fauna estimated to be of lower oligocene age (for an example see the crab in fig. 6e). of probable miocene age is the very restricted kap brewster formation, which has a basal breccia suggesting proximity to a major fault and thus providing evidence of tectonic movements sometime between the eocene and miocene. larsen, m. et al. (2005) carried out a rather exhaustive study which includes details of earlier work. they concluded that the kap dalton formation was formed in a limited period in the middle eocene (chrons 21r to 20r) during transgression over a muddy shelf with two distinctive quartzitic sandstone units of unknown, but distal, provenance. dating has been refined by palynological studies by jolley (1998) to early eocene (52.5 ma) for the bopladsdalen formation and middle eocene (c. 43 ma) for the krabbedalen formation. an early eocene age is different from the radiometric ages of the uppermost main basalts (fig. 3), although the discrepancy, if not due to error, may be explained by longer-lasting volcanism in this coastal area. miocene basalts as a result of detailed field work in 1995, very young lavas were unexpectedly found in a small area to the north of the borgtinderne intrusion (fig. 2) overlying the plateau lavas and occurring as remnants capping nunataks. named the vindtoppen formation, they have given 14–13 ma ages (miocene) by the 40ar/39ar method (storey et al. 2004). they are transitional to mildly alkaline with isotopic evidence of contamination by archaean material. melting of metasomatised mantle during regional uplift was tentatively suggested as the cause of this volcanism, but its relations to the overall development of the margin remain obscure. 29 offshore basalts seaward-dipping reflectors (hinz et al. 1987) were first imaged seismically along many continental margins and interpreted as tilted lava flows – an interpretation later confirmed by drilling. seaward-dipping reflectors are characteristic of so-called volcanic margins, of which east greenland is a prime example (larsen, h.c. & jakobsdóttir 1988): they are present over most of the margin. in addition, large areas of the shelf are covered with basalt extensions from onshore (fig. 1). the ocean drilling program drilled the basaltic basement on legs 152 and 163 (larsen, h.c. & saunders 1998), sampling such reflectors. locations of these sites are shown in fig. 1, with shelf sites 915, 917 and 918 located on the landward feather-edge of the seawarddipping reflectors. at site 917, 779 m of basalt were penetrated, and a sequence of steeply dipping, metamorphosed sandstones and siltstones was reached (vallier et al. 1998). these sedimentary rocks are of unknown age, but lithologically they are equated to the onshore ryberg formation. the volcanic stratigraphy was divided into three units: a lower series of continentally contaminated basalts with rare picrites, a middle series of more evolved, contaminated basalts, dacites and acid tuffs and, following a sandstone layer, an upper series of basalts and picrites of mid-ocean ridge affinity. at site 915, only a few metres of basalt were penetrated. at site 918, 235 m of the basaltic basement to the sedimentary rocks were sampled; a sill or flow within early eocene sedimentary rocks lies 9 m above the basaltic basement. werner et al. (1998) reported polymict breccias and conglomerates overlying the basalts at sites 917 and 918, perhaps equivalent to the hængefjeldet formation. tuffs and lapillistones within the basalt formation yielded ar-ar ages of 60.5 ± 0.2 and 60.2 ± 0.2 ma. on leg 163 (larsen, h.c. et al. 1999), holes at sites 989 and 990 supplemented the earlier holes and penetrated the basalts. they showed oceanic affinities similar to the upper series at sites 915 and 917. unfortunately, leg 163 was terminated early due to the ship being hit by a hurricane (odp shipboard scientific party 1996) and did not achieve its goals (which were: to complete the leg 152 traverse and to drill an additional traverse closer to the plume source). to summarise the offshore results very briefly: the recovered spectrum of compositions span a very wide range – approximately covering that of the entire north atlantic igneous province. there is abundant evidence for the involvement of continental materials, not least the presence of acid rocks similar to those found on the eastern atlantic margins (see sinton et al. 1998 and references therein). as regards the oceanic rocks in the upper part of the sequence, there is some evidence that they formed in a cooler mantle than those closer to the supposed plume centre. reliable radiometric ages are in the range 62 to 60 ma (sinton et al. 1998; werner et al. 1998). these lavas are subaerial and therefore most similar to the lower basalts. the sub-basaltic sedimentary rocks may be equivalent to the kangerlussuaq group. however, the onshore sedimentary rocks do not show the same high-angle unconformity with the overlying basalts and are unmetamorphosed. although dacites and acid tuffs have not been found in the lower basalts, these are strongly contaminated with continental material (hansen & nielsen 1999). on the basis of existing evidence, it is possible that thick dacite sequences only occur peripherally to the postulated plume head. tegner & duncan (1999) summarised results from the two east greenland legs of the ocean drilling program, which span c. 12 ma and fall into three mantle melting events. pre-break-up continental volcanism occurred at c. 61–60 ma, followed by syn-break-up volcanism c. 57 ma and post-break-up volcanism c. 49.6 ma. this was interpreted as heralding the arrival of the plume head under central greenland, continental break-up and finally the passage of the plume stem beneath the east greenland margin. basalts of northern east greenland scattered outcrops of a palaeogene lava succession occur between 73 and 76°n along the east greenland margin, notably at hold with hope, wollaston forland and the island of shannon, with a small outcrop on traill ø (fig. 8). thick sills are also abundant in the sedimentary basins, both in the basalt areas and to the south, as in the jameson land basin (larsen, h.c. & marcussen 1992). the lavas, with extents shown in figs 2 and 8, have been described in a number of papers (noe-nygaard & pedersen 1974a, b; hald 1978, 1996; upton et al. 1980; upton et al. 1984a; upton 1988; upton et al.1995; ellam et al. 1998). the same area also includes the myggbukta complex, a subvolcanic caldera complex, generally intensely hydrothermally altered (upton et al. 1984b). farther east, at kap broer ruys, there is another probable central complex which 3030 is not exposed apart from some minor granophyric and felsitic bodies. the lavas are divided into two series: a lower series of c. 450 m and an upper series, c. 350 m thick. the lower series is quite uniform, aphyric to microphyric and consists of quartz tholeiites. upton et al. (1995) suggested that these had been erupted far to the east, on the present vøring plateau off norway during a massive burst of volcanic activity connected with the arrival of the icelandic plume and continental breakup. trace elements and isotopic ratios (thirlwall et al. 1994) show little evidence of crustal contamination and there are close similarities to icelandic basalts, as well as those of the vøring plateau. the upper series is much more variable, consisting of olivine tholeiites, alkali basalts and some basanites: they are commonly highly porphyritic and show evidence of interactions with subcontinental lithosphere. a much greater proportion of these magmas relative to the lower series was derived from garnet-facies mantle. these lavas are also thought to have been derived from a more local source than those of the lower series. 40ar/39ar dates on two dykes related to the upper series give values of c. 56.7 and 56.6 ma, and an ankaramite flow gave 53 ma, although this was thought to have suffered ar loss. a basal nephelinite unit gave 58.7 ma, but the possibility that this might be a sill cannot be ruled out (b. upton, personal communication 1998). the finding of elevated 3he/4he ratios (total range from 1.7ra to 21ra, where ra is the atmospheric ratio) in some of these basalts is fully consistent with present interpretations of origin from a mantle plume (marty et al. 1998). similarly high ratios were reported by bernstein et al. (1998b) for intrusions south of scoresby sund. loch fyne hold with hope 75° 74° giesecke bjerge 24˚ kap pansch myggbukta clavering ø kejser franz joseph fjord kap broer ruys jackson ø finsch øer eskimonæs kap stosch wollaston forland kap philip broke hochstetter forland alabama shannon lille pendulum bass rock moskusoksefjord kuhn ø 75˚ 74˚ 20˚ young sund germania havn tyro ler fjord b ad landsdal kap david gray h ochstetterbugten sabine ø dronning augusta dal gauss halvø fligely fjord plutonic centres basaltic lavas and sills basement 50 km fig. 8. approximate extent of basic rocks in northern east greenland, mainly basalts although some are extensive sills. map compiled from references cited in the text. two small, central complexes (upton et al. 1984b) are also shown. location shown in fig. 2. 31 most of these lavas are roughly contemporaneous with the bulk of the lavas in the kangerlussuaq – scoresby sund area to the south and were erupted during magnetochron c24r, although part of the upper series may have erupted during the succeeding normal and reversed epochs. the most reliable age for the latest dyke swarm comes from a picritic dolerite sheet and is c. 33 ma. long-distance tephra falls the sanidine-bearing tuff in the uppermost part of the skrænterne formation was mentioned above and will also be discussed with respect to its age in the section below on nephelinites. heister et al. (2001) argued, on the basis of its age, chemistry and highly distinctive mineralogy, that it had originated in the gardiner complex. they further identified this ash in the diatomaceous fur formation in northern denmark (in the limfjorden area where it is well exposed on the islands of fur and mors; brooks 2006) as the layer designated –17. this is also of a distinctive alkaline composition, has identical rare-earth element patterns, contains unusual minerals like aegirine, katophorite and perovskite and is of the right age. the same ash occurs at several sites in the north sea (e.g. knox 1996) and was identified in ocean drilling program site 550 on goban spur, south-west of ireland. if the identification of heister et al. (2001) is correct, it makes this tephra a very important stratigraphic marker. this is especially so as the eruption can be related to the paleocene–eocene thermal maximum, a short-lived warming event which has attracted much interest, especially as to its cause. the sanidine-bearing ash, which is designated as layer –17, has been dated by a number of workers using the 40ar/39ar method (see westerhold et al. 2009 for a summary) in an attempt to accurately date the thermal maximum and to refine the geological time scale, including the cretaceous–paleocene boundary. astronomical tuning (which involves identifying and counting milankovich cycles, whose durations are known) in sedimentary rocks of this part of the succession indicates that the ash of layer –17 post-dates the thermal maximum by 672 ± 21 ka. the fur formation is characterised by c. 180 ash layers, mostly of a high-titanium ferrobasalt type, similar to many in east greenland and the faroe islands. all these ashes have also been studied in detail by larsen, l.m. et al. (2003), who defined their compositions and attempted to deduce their source areas. some other layers (particularly those designated –11, –12 and –13) are similar to the alkali basalts in the upper part of the east greenland succession. larsen, l.m. et al. (2003) concluded that the lower fur ashes stem from the general east greenland area, but later, more voluminous ferrobasaltic ashes originated from the already submerged, evolving rift between east greenland and the faroe–rockall area: the site of the postulated icelandic protoplume. the correlation of layer –17 and nearby highly alkaline ashes with the gardiner complex was only tentatively accepted. another prominent ash layer in the danish succession is labelled +19 and is also a potentially important stratigraphic marker. this is a thick, peralkaline rhyolitic ash, postdating the paleocene–eocene thermal maximum by 862 ± 21 ka according to astronomical tuning (westerhold et al. 2009). the source of this ash could not be identified by larsen, l.m. et al. (2003), but similar compositions are found in east greenland and, indeed, in iceland, although the icelandic ones are much younger and cannot be a source. this ash, which was also identified in austria by egger & brückl (2006), was estimated to have a total volume of 1200 km3, placing it among the largest volcanic events known. some basaltic ash layers were estimated to have volumes of c. 500 km3, whereas many are c. 120 km3, which makes them similar in volume to the largest historical eruption (i.e. tambora). it should be noted that the lower part of the danish ash series is dated by the –17 ash to be contemporaneous with the uppermost skrænterne formation, so the bulk of the ashes (in the positive series, mostly basaltic) must have been erupted after the east greenland main basalts. likewise, they cannot have come from the faroe islands or britain as volcanic activity in these areas had ceased at this time (at least, on the basis of the presently known geology). the conclusion is that these eruptions took place after continental break-up, in areas now located offshore, in agreement with the conclusions of larsen, l.m. et al. (2003) as described above. they record well over 100 truly vast eruptions. egger & brückl (2006) estimated that the 23 ash layers they identified in austria represent a total eruption volume of c. 5000 km3. 3232 gabbroic intrusions the skaergaard intrusion although only a superficial description of research on the skaergaard intrusion is given here, a brief overview is appropriate as many of the problems encountered in this intrusion also apply to the other gabbros of east greenland. the history of research on the intrusion also provides an interesting insight into changing fashions, not to mention the difficulty in reaching a consensus on even the most basic questions in igneous petrology. the intrusion was discovered by l.r. wager on his first expedition to greenland in 1930–1931, known as the british arctic air route expedition (watkins 1932a, b, c; brooks 1985b, 1990; hargreaves 1991; glasby 2007a, b). he not only recognised the scientific significance of the intrusion, but had the tenacity to overcome the immense difficulties of travel to east greenland at that time, returning to overwinter in 1935–1936 (wager 1937) and spending the year exploring a vast area by dog-sledge. the early years of research on the skaergaard intrusion were taken up with the basic description of its structure, mineralogy and composition as published by wager & deer (1939). in the 1950s, the thrust of wager’s research was the mechanisms of the cumulus process, and a new visit to the intrusion took place in 1953. in the 1960s, emphasis moved to geochemistry as sensitive techniques for the determination of trace elements became available, and the distributions of a large number of elements were documented, largely by oxford dphil students, and the principles of their distribution in magmatic rocks determined. no scientific visits to the intrusion took place in the period 1953–1970 when only prospectors visited the area. in 1971, groups from the university of oregon and the university of copenhagen were present, and a new research impetus emerged such that over the following years, researchers from several universities and other research institutions, particularly stanford university, worked in the field. the main research during this time was in examining the cumulus theory of wager, proposing other mechanisms, and the study of the hydrothermal systems of the intrusion. amazingly, it was not until 55 years after its discovery that the intrusion (and others nearby) was found to contain a world-class precious metal deposit (bird et al. 1991, 1995; arnason et al. 1997a, b; arnason & bird 2000; nielsen et al. 2005). in 1986, the canadian company platinova resources limited arrived and rapidly found gold anomalies in the skaergaard intrusion, sparking a new phase of diamond drilling and the study of precious metals. the danish lithosphere centre initially targeted the volcanic succession, and it was not until the year 2000 that samples from the intrusions were collected. in that year, new comprehensive collections from the skaergaard, kangerdlugssuaq and other intrusions were made (nielsen et al. 2001). in recent years, samples have also become available from commercial drill cores as well as from the 2000 collection. work has been directed to revising the structure of the intrusion (nielsen 2004), mainly using drill-core material. subsequently, microtextural studies have been employed, with the aim of getting a better understanding of the processes taking place during the crystallisation of such rocks, by groups centered at cambridge university (e.g. holness et al. 2006, 2007, 2011; humphreys 2009, 2011; humphreys & holness 2010; mckenzie 2011) and at the university of aarhus (tegner et al. 2009). the skaergaard intrusion is not a large intrusion. it measures 11 km north to south, 8 km east to west, and is 3.4–4 km thick. its volume is estimated to be 280 ± 23 km3 (nielsen 2004). originally, wager & deer (1938) regarded its shape as being that of a funnel, but several alternative models have been proposed: laccolith, box-like, balloon-like (see irvine 1992; nielsen et al. 2005); no consensus seems to have emerged. however, it is agreed that the intrusion was filled fairly rapidly, and no further pulses of magma were added or extruded, so that differentiation took place in a closed system of a fairly normal type of tholeiitic magma. the internal divisions of the skaergaard intrusion are well known: the marginal and upper border series where magma has crystallised adjacent to the walls and roof, respectively, and the layered series. according to wager & deer (1939), crystals in the latter series accumulated under the influence of gravity, leaving a clear record of changing magma compositions and forming gravity-generated structures such as layering with modal sorting, slumping cross-bedding and other features. some of these are enigmatic, e.g. christmas-tree-like dendrites (fig. 11g) and the trough bands investigated in depth by irvine & stoesser (1978). unfortunately, the promise of tracking magma variations from cumulate rock and mineral compositions turned out to be utopian. at least two views have 33 emerged: the so-called bowen trend, in which differentiation leads to silicic products, and the fenner trend, where there is increasing iron concentration (see the controversy: hunter & sparks 1987; brooks & nielsen 1990; mcbirney & naslund 1990; morse 1990). the difficulty in determining the line of liquid descent is, of course, that we have no liquids and must extrapolate from the rocks. a further process that may be important in the development of basaltic magmas is liquid immiscibility, which was out of favour for many years but has now received renewed credibility (mcbirney & nakamura 1974; naslund 1980; jakobsen et al. 2005, 2010; humphreys 2011). general agreement on these problems has not yet been reached. central to l.r. wager’s view of the intrusion was the sinking of crystals (primocrysts) in the magma under the influence of gravity, first adopted by wager & deer (1939) and later treated in detail by wager et al. (1960) and in the book by wager & brown (1968). this led to a completely new nomenclature as it was claimed that the primocrysts (cumulus crystals) in the rock could be distinguished from the intercumulus material, or trapped liquid. various proportions of intercumulus material could be distinguished such that orthocumulates were rocks in which the cumulus crystals were from the instant of sedimentation, and the interstitial material trapped between the crystals would be equivalent to the magmatic composition at the time of deposition. at the other extreme, the intercumulus liquid had exchanged with the magma, and the primocrysts had overgrown to eliminate all the interstitial liquid, forming adcumulates. mesocumulates lay in between these two end-members. this entire view of cumulate rocks was later attacked, originally by mcbirney & noyes (1979), but later also by other workers, who claimed that the primocrysts had grown in place against the walls and roof in boundary layers with strong concentration and temperature gradients where double diffusive convection was operative, and the magma chamber could split into separately convecting layers. this is an issue which still divides igneous petrologists, although it seems to have faded into the background in recent years. irvine et al. (1998) showed that many structures in the skaergaard layered series are undoubtedly gravity-derived and contend that crystals did sink, and there were convection currents. important studies were also conducted at the skaergaard intrusion on the circulation of meteoric waters, beginning with the studies of norton & taylor (1979), taylor & forester (1979) and norton et al. (1984), and continuing with a large number of papers by a group from stanford university (bird et al. 1986, 1988; manning et al. 1986, 1993, 1995; manning & bird 1991). these studies extended to other nearby gabbros (e.g. rose & bird 1987, 1984; fehlhaber & bird 1991). norton & taylor (1979) were able to present a detailed cooling model for the intrusion, which could be refined when accurate ages were obtained for minerals with different blocking temperatures. for a long time, the age of the skaergaard intrusion was poorly known. a fission track date indicated an age of 54.6 ± 1.7 ma (brooks & gleadow 1977), but this was not widely accepted due to the low repute of the method with traditional geochronologists. subsequent 40ar-39ar ages on biotite were published by hirschmann et al. (1997) giving 55.48 ± 0.30 and 55.40 ± 0.14 ma (plateau and isochron ages, respectively). finally, u-pb zircon dating of the sandwich horizon rock (supposed to be the latest differentiate, sandwiched between the layered series and the upper border series) gave 55.59 ± 0.13 ma (hamilton & brooks 2004), and very high precision ages have recently been published for the sandwich horizon and the overlying basistoppen sill (wotzlaw et al. 2011). despite more or less continuous studies over more than 70 years, some of the fundamental questions related to such intrusions and igneous processes in general (e.g. the nature of the liquid line of descent, the importance of immiscibility (mcbirney & nakamura 1974); the causes of igneous layering, the origin of precious metal concentrations) have not yet been answered satisfactorily (see e.g. wilson 1993; young 2003). indeed, the concept of magma chambers is doubted by some; none of any size having been imaged seismically anywhere in the world (dettrick et al. 1987; lundstrom 2009). the skaergaard intrusion has many of the features shown by other gabbros in east greenland (see below), but differs in some respects. it is the only body which crystallised as a closed system, i.e. the magma chamber was filled by a single pulse, and no extrusions of lava took place to the best of our present knowledge. also, the magma developed on an unusual iron-rich trend, probably as a result of the closed system evolution. finally, there are no ultramafic rocks as found in several other intrusions nearby (the so-called wehrlite suite). 3434 other gabbroic intrusions in addition to the skaergaard intrusion, gabbroic intrusions are ubiquitous throughout the area (fig. 2). notable examples are: kap edvard holm (elsdon 1969, 1971a, b, 1972, 1982; abbott & deer 1972; fehlhaber & bird 1991; bernstein et al. 1992; tegner & wilson 1993, 1995; tegner et al. 1993; brandriss et al. 1995; bernstein et al. 1996; brandriss et al. 1996; tegner et al. 1998; brandriss & bird 1999), kærven (ohja 1966), kruuse fjord (arnason et al. 1997a, b; tegner et al. 1998b), imilik (brown & farmer 1972; tegner et al. 1998b) and the newly discovered noe-nygaard intrusion (bernstein & bird 2000). maps of the kap edvard holm and imilik gabbros are shown in fig. 9. these are probably the two largest onshore gabbros in east greenland, but both are imperfectly known. they are both made up of three main units of gabbro and include occurrences of wehrlite. several offshore gabbros are exposed only on small islands: nordre aputiteeq, ittutarajik, pattuulaajivit, søndre aputiteeq and an unnamed group of rocks informally known as ‘the skerries’. their underwater extent, like that of imilik, is unknown. these island gabbros give an insight into the nature of what must be very extensive gabbros on the shelf and are of considerable interest themselves. some information has been published by bernstein et al. (1998b) and tegner et al. (2008), although many unpublished observations exist (e.g. petersen 1977). hitherto unpublished maps by d.k. bird and n. rose are given in fig. 10. the majority of these gabbros are tholeiitic. the gabbros of the werner bjerge complex and its satellites to the north (see below) are of alkaline character (bearth 1959; kapp 1960; brooks et al. 1982), but there are w w w w w 34°w agtertiat upper layered series bom bardier gletscher lower layered series kontaktbjerg stribede nunatak boswell bjerg ? ? 85 80 55 65 80 80 70 70 80 70 80 70 80 80 80 80 85 78 55 80 85 70 53 40 50 imilik ø kap s.m. jørgensen taco pynt kap boswell barberkniven den lave pynt kap deichmann kap edvard holm keglen polaric g letscher hutchinson gletscher sdr. b osw ell bugt poulsen fjord 66 5 kmarchaean gneiss basalt gabbro syenite intrusion breccia granite w archaean gneiss2.5 km pre-rift basalts gabbro phase i gabbro phase ii gabbro phase iii wehrlite intrusions 32°w 68°n middle layered series 66°45’n a b fig. 9. maps of two major intrusions. for locations see fig. 2. a: kap edvard holm complex. map modified from abbott & deer (1972). geochronological data suggest a value of c. 49 ma for the lower layered series and about 1 ma less for the middle layered series. the satellite syenitic intrusions were described by deer et al. (1984). b: imilik intrusion (compiled from unpublished maps by s. bernstein, d.k. bird and c. tegner). 35 no stable isotope studies from this northern area. gabbro also occurs in the myggbukta area of hold with hope. the timing of its emplacement has been discussed by upton et al. (1984b). the lilloise intrusion (fig. 2) is reported to have been derived from an alkali basalt magma (brown 1973). it is associated with large amounts of ultramafic rocks and underto oversaturated minor syenitic bodies (brown 1973; matthews 1976; brown et al. 1987; chambers & brown 1995; bernstein et al. 1998b; tegner et al. 1998b). recently, a new gabbro was described from the kialeeq district (bernstein & bird 2000), and named the noe-nygaard gabbro (see below). this intrusion contains wehrlites, as described previously for other gabbros of the area, which the authors explained as being generated by the dehydration of stoped, altered basalt xenoliths in accordance with previous studies. the seminal oxygen isotope studies of norton & taylor (1979), taylor & forester (1979) and norton et al. (1984) on the skaergaard intrusion have already been mentioned, and a study of hydrogen isotopes in the gabbros and adjacent syenites of kap edvard holm, skaergaard and nordre aputiteeq led nevle et al. (1994) to suggest that at between 55 and 43 ma these gabbros equilibrated with meteoric water flowing from the high altitudes existing inland at that time, due to the dome described by brooks (1979). sheppard et al. (1977) showed that, unlike all other studied gabbros of the north atlantic igneous province, the lilloise intrusion does not record the influx of meteoric water. instead, the aureole has been altered by magmatic water flowing from the intrusion. while documentation of all these bodies is far from complete, it appears that most are later than the skaergaard intrusion and, in contrast to the latter, apparently display multiple fillings of magma chambers, such as has been clearly documented in the lower layered series of the kap edvard holm complex (bernstein et al. 1992) or even multiple intrusions (e.g. imilik as shown by tegner et al. 1998b and kap edvard holm). they appear to have been intruded in several time windows (tegner et al. 1998b): 56–54 ma, coeval with the bulk of the basalts (the skaergaard intrusion, part of the imilik intrusion and the sorgenfri gletscher sills – see below), 50–47 ma (most of the remaining) and 37–35 ma (mostly in the mesters vig and kialeeq (kialineq) areas – see below), which, as discussed by tegner et al. (1998b, 2008), may also be the age of some of the latest lavas (with the exception of the vindtoppen formation). the oldest unit at imilik (tegner et al. 2008) has not been dated but is believed to have an age of c. 60 ma and thus coeval with the lower basalts and the earliest phase of activity throughout the north atlantic igneous province. the intermediate unit is >56.6 ± 0.6 ma (age of a cross-cutting dyke), and the youngest unit has given 40ar/39ar ages of 50.1 ± 1.2 and 49.5 ± 0.2 ma. it will be shown that these three ages coincide with main periods of magmatism in east greenland, and the late age of many of the east greenland gabbroic intrusions may reflect the passage of the east greenland margin over the mantle plume stem, a suggestion also made by bernstein et al. (1998b) based on rare-earth elements, sr, nd and he isotopes. holm et al. (2006) published a zircon date for the kærven gabbro, which is not completely concordant with the date provided by tegner et al. (2008). however, both are considerably younger than the age of c. 58 ma inferred by holm (1991), and the kærven intrusion falls into the 56–54 ma window. the picture that has emerged of the timing of gabbro intrusions (and by extension probably also of surface volcanism) is a series of pulses at approximately 60, 55, 49 and 35 ma. these pulses seem to be separated by intervals of c. 5 ma, except that a pulse at c. 40 ma is absent (but may be represented by some of the mesters vig gabbros). in iceland, major rifting events involving rift jumps occur on an 8 ma time frame (hardarson et al. 1997). an interesting study of a gabbro plug at courtauld fjord (bernstein 2006) included spot analyses of pyroxene crystals using electron microprobe and laser ablation icp-ms methods. they showed that the compositions of the parent magmas (a long-standing problem when dealing with plutonic rocks) could be estimated using partition coefficients for many elements. the lilloise intrusion and its near surroundings were used as an example in a book by mosely (1981) on field methods. the relevant chapter includes maps, photographs and methods used in the field, a discussion of the glacial deposits and a section on the basalts and dykes. spectacular layered features in the skaergaard intrusion are known world-wide, and similar features occur in most of these gabbros. figure 11 shows a selection of some typical structures encountered in east greenland gabbros. these include various types of layering, such as macroand inch-scale, wispy and cross-bedding (fig. 11a), rhythmic, apparently gravity-controlled, slumped layers (fig. 11b), layers deformed by settling autoliths (fig. 11c) or rising diapirs (fig. 11d, e) etc. in addition, 3636 many other features, some perplexing, may be present. these include anorthositic autoliths or xenoliths and cross-cutting anorthosites (fig. 11e, f), christmas-treelike and other dendrites formed by plagioclase in the skaergaard intrusion (fig. 11g) formed by olivine (over 30 cm long) at kap edvard holm (tegner et al. 1998b) and by pyroxene and plagioclase in the miki fjord macrodyke (blichert-toft et al. 1992). also areas of apparent pyroxenite replacement and other features have been described (e.g. mcbirney & sonnenthal 1990; sonnenthal 1992). the exposures on ice-polished surfaces are often exceptionally fine, making these rocks worthy and rewarding objects of study. an additional impetus for research on the basic intrusions came with the discovery of very large and potentially economic amounts of precious metals in the skaergaard intrusion (bird et al. 1991; andersen et al. 1997; nielsen et al. 2005), closely followed by reports of precious metal occurrences in neighbouring gabbros (e.g. bird et al. 1995; arnason et al. 1997a; arnason & bird 2000,) and the miki fjord macrodyke (discussed below). these discoveries led to extensive diamond drilling, and the cores are at present being studied from a scientific point of view (nielsen et al. 2000). a detailed description of the skaergaard deposit is given by nielsen et al. (2005). it is known to be a world-class deposit containing 10 million ounces of gold and 30 million ounces of palladium (platina resources limited: www.platinaresources.com.au), and exploration work including diamond drilling continues at the time of writing (2011). both these metals are highly valued at the time of writing and unlikely to lose value in the near future. the juxtaposition of gabbros (with multiply refilled magma chambers and a late ultramafic suite – see next section), intense dyke swarms and basaltic lavas is reminiscent of ophiolite complexes, except for the presence of continental rocks and lack of a mantle unit or pelagic sediments. the association represents the earliest stages of ocean-floor formation and to this extent may be regarded as a continental margin ophiolite (bernstein et al. 1992). a major problem with the gabbros is their mode of emplacement, classically known as the ‘room problem’. in no case is this very clear, and as stated above the emplacement mechanism of the skaergaard intrusion, the most studied of all, is still in dispute. ultramafic plugs etc. in the area to the immediate north and east of kærven, several plug-shaped ultramafic bodies occur. they have nordre aputitêq (nordre aputiteeq) pátûlâjaivit (pattuulaajivit) igtutârajik (ittutarajik) 67°45’n 3 2 °w deception ø basalt gabbro gabbro kap edvard holm fladø ‘the skerries’ basalt gabbro 5 km basalt basement dip > 5% d yk es < 5% d yk es d a n m a r k s t r æ d e > 50 % d yk es tuff 40 30 30 fig. 10. maps of gabbroic intrusions. for locations see fig. 2. left: overview of intrusions on the islands and skerries to the south of kap edvard holm, suggesting the extent of offshore gabbro. facing page: gabbros of the individual small islands south of kangerlussuaq with ages of c. 50–48 ma (tegner et al. 1998b), compiled from unpublished maps by d.k. bird and n. rose (1987–1990). see caption to fig. 2 and names in brackets for modern geographical spellings. 37 85 30 35 35 inuit house ruin hydrothermal veins mafic dykes orientation of lavas orientation of lavas dip of layering in gabbros no. of veins 40 20 0 20 microgranite-gabbro breccia zone of metabasalt and metasediment blocks gabbro diorite granite 3 2 °2 6 ’ 3 3 °2 0 ’ 3 2 °3 0 ’ 60 35 80 45 45 50 55 80 80 80 74 60 55 55 gabbroic dykes dip of layering 43 67°48’ 67°37’ c en tr al co m pl ex m ar gi na l g ab br os 35 50 40 40 30 30 35 35 45 30 30 30 30 40 40 40 32 32 40 31 38 45 38 38 44 30 30 weather station (abandoned) oliebugt yankeetoppen gabbro (undifferentiated) snow and ice basalt and tuff ‘the skerries ’ 3 2 °1 5 ’ 3 2 °3 8 ’ gabbro base of intrusive unit mafic plugs metabasalt xenoliths snow and ice nordre aputitêq basalt and tuff gabbro kaersutite gabbro granite snow and ice n30 85 veins per metre at contact n igtutâraj ik søndre aputitêq 500 m pátûlâj iv it 500 m 67°47’ 67°15’ 67°36’ snow and ice 500 m 500 m fugleø inuit house ruin n n 500 m ( i t tutar a j i k )(nordre aput i teeq) (søndre aput i teeq) (pattuu laa j iv i t ) mafic dykes snow and ice volcanic rocks gabbro faults 35 orientation of lavas 80 orientation of faults n n 3838 a c b d g e f h 39 been studied by prægel & holm (2001) and holm & prægel (2006). in general, rock types from these intrusions include: dunite, wehrlite, lherzolite and olivine clinopyroxenite. the above-mentioned authors concluded they are cumulates from a magma of icelandic type which was contaminated by local archaean continental lithospheric mantle. some intrusions are cut off by the kærven and kangerdlugssuaq intrusions, whereas others are clearly seen as plug-shaped, e.g. an almost circular body several hundred metres in diameter just south of the snout of the glacier that descends from kangerlussuaq tinde (fig. 11h). it was called mordor by holm & prægel (2006), who reported that it is zoned with an ultramafic outer part and a gabbroic core. a similar wehrlitic body was discovered just north of the contact of the skaergaard intrusion at watkins fjord as a result of gravimetric studies of the skaergaard intrusion (blank & gettings 1973). it is poorly exposed, but by comparison with the ultramafic rocks on the western side of the fjord, it is interpreted as a plug. indeed, the largely moraine-covered peninsula on which this body occurs is almost circular, suggesting that the wehrlite determines the shape of the peninsula. this body was mapped by mcbirney (1997) as precambrian, but it has palaeogene isotopic characteristics (stewart & depaolo 1990; s. bernstein, personal communication 1995). as it may well be intersected by the skaergaard intrusion at depth, it may be the source of the so-called gabbro picrite inclusions in the northern marginal border series of the skaergaard intrusion, which wager & deer (1939) considered to be early cumulates rafted up from the hidden layered series. the first wehrlites from the kap edvard holm complex were reported by bernstein et al. (1992, 1996) and tegner and wilson (1993). they are known as the wehrlite suite, although not all are strictly wehrlites, but peridotites sensu lato. they are mica-bearing and both intrude and replace the gabbros; fig. 11d shows anorthosite diapirs emplaced into a wehrlite. these rocks are strongly reminiscent of ultramafic rocks in ophiolite complexes (juteau et al. 1988). since the initial discoveries, ultramafic components have been reported from many layered gabbros in east greenland (brandriss et al. 1995, 1996; brandriss & bird 1999; bernstein & bird 2000). it has been suggested that such ultramafic magmas form by suppression of plagioclase crystallisation caused by the release of volatiles from hydrothermally altered basaltic xenoliths, a process documented in some detail in the newly discovered noe-nygaard intrusion (bernstein & bird 2000). interestingly, the skaergaard intrusion lacks this suite of rocks, possibly because it does not have a large volume of basaltic xenoliths (irvine et al. 1998). a remarkable dyke in the astrophyllite bay complex at astrofyllit bugt in amdrup fjord (fig. 12) is crowded with angular, ultramafic xenoliths (brooks & nielsen 1982a). it cuts a large gneiss inclusion and other field relations cannot be determined, except that it is older than the syenite (and younger than the gneiss). it yields a k-ar age of c. 52 ma, but the isotopic systematics may have been reset by the intrusion of the syenite. ultramafic inclusions (largely harzburgites) in dykes at wiedemann fjord will be discussed in the subsequent section on dykes and sills. felsic intrusions and volcanic rocks an important characteristic of the east greenland rifted margin is the large amount of felsic intrusive rocks of broadly syenitic character (figs 12, 13). the distribution of syenites in the kangerlussuaq area is shown in fig. 12, in the mesters vig area in fig. 14 and in the kialeeq (kialineq) area in fig. 16. figures 13, 15 and 17 depict the occurrence and textures of the syenites in these three areas. felsic, volcanic rocks are almost facing page: fig. 11. magmatic structures in gabbros. a: discordant layering on the island of pattuulaajivit. b: folded and disrupted layering in the kap edvard holm lower layered series. c: face of ‘gabbrofjeld’ (about 500 m high), skaergaard intrusion, showing various layered features including macrorythmic layering (lower part of picture), draping a large autolith (lower centre) and the triple group (three-quarters up from the bottom of the picture). the latter is associated with precious metal mineralisation. d: anorthositic diapirs intrusive into a wehrlite, kap edvard holm lower layered series. e: anorthositic diapir in the lower zone of the skaergaard intrusion. f: anorthositic xenoliths/autoliths in the kap edvard holm middle layered series at den lave pynt (fig. 9a). g: dendritic plagioclase lying in the plane of layering in the lowest unit of the upper zone (uza) of the skaergaard intrusion (sonnenthal 1992). h: ultramafic plug (arrow, ‘mordor’ of holm & prægel 2006) cutting the basement on the west side of kangerlussuaq. the high, snow-capped peak in the background is kangerlussuaq tinde (c. 2200 m). 4040 unknown except for a small outcrop in kialeeq (brown & becker 1986) and various porphyries infilling the calderas of kap parry and kap simpson (figs 14a, 15b; schaub 1938, 1942). in the kangerlussuaq area, brooks & nielsen (1982a) showed that, at the present level of exposure, syenite occupies approximately twice the area of gabbro. felsic intrusions granites, syenites and nepheline syenites mafic and ultramafic intrusions mainly gabbro, minor peridotite volcanic rocks mainly tholeiitic plateau basalt sedimentary rocks kangerdlugssuaq group, cretaceous–eocene precambrian gneisses largely archaean 10 km 2660 m domkirkebjerget k angerlussuaq f r e d e r i k s b o r g g l e t s c h e r watkins fjord amdrup fjord bagnæsset kræmer ø astrofyllit bugt flammefjeld trebjørnebjerget citadellet skærgården skaergaard intrusion kap hammer kap deichmann barberkniven kap boswell kap edvard holm 68°30'n 68°n 68°30'n 31° w32°w 33°w 32° w m ik i fjord batbjerg amdrup pynt polaric g letscher sdr. syenit gletscher nordre aputiteeq kangerlussuaq gletscher kangerdlugssuaq intrusion gardiner intrusion batbjerg complex (caledonian) kap edvard holm intrusion k æ lv eg le ts ch er kærven n ordfjord g letscher sødalen i.c.jacobsen fjord hutchinson gletscher kangerlussuaq tinde 2200 m c ou rt au ld g le ts ch er fladø fig. 12. geological map of the kangerlussuaq area (based on publications of l.r. wager and the author’s observations). location shown in fig. 2. 41 the kangerlussuaq – blosseville kyst area the largest of the east greenland syenites is the kangerdlugssuaq alkaline intrusion (its name is the old spelling of the fjord kangerlussuaq) and its satellites (figs 12, 13a, b, c), covering >800 km2, making it among the largest syenitic bodies in the world. it has been described in detail by wager (1965), kempe et al. (1970) and kempe & deer (1970, 1976), and its satellite intrusions by deer & kempe (1976). broadly, it zones from oversaturated, almost granitic rocks (‘nordmarkites’, fig. 13c) adjacent to the country gneisses over syenites (known as pulaskites in the literature) to strongly undersaturated nepheline-sodalite syenites (‘foyaites’, fig. 13b) at the core. these plutonic rocks are usually spectacular and massive, with very few joints. some are strongly laminated, others have eye-catching dark, perthite phenocrysts several centimetres across (– feldspars were studied in some detail by kempe 1966). brooks & gill (1982) reported additional information on the mineralogy of the intrusion and favoured the idea that the most primitive magma was represented by the undersaturated core (a reversal of all previous thinking), and that these magmas had reacted with granitic country rock leading to the quartz-oversaturated peripheral rocks; a scenario consistent with the available isotopic information (pankhurst et al. 1976). other workers (wager 1965; kempe et al. 1970) have considered the progression of rocks to be the reverse and the main agent to have been crystal fractionation. however, they were not able to explain the mechanism by which the thermal barrier in the residual system (quartz-nepheline-kalsilite) has been crossed. riishuus et al. (2006, 2008) reported new geochemical and isotopic data from the kangerdlugssuaq intrusion with determinations of major elements, supplied a comprehensive list of trace elements and sr-nd-hf-pb isotopes, using a new sample collection. they essentially agreed with the earlier model of brooks & gill (1982) in that differentiates of mantle-derived basic magmas had reacted with the gneissic basement in deep mantle chambers, where periodic injections into the main chamber became progressively less contaminated. riishuus et al. (2005) also reported on the nearby astrophyllite bay complex where basic pillows were previously claimed by nielsen & brooks (1981) to have caused melting of the gneiss, and diffusive reactions had formed the enclosing syenite. the new study shows this not to be so: the syenite and the basaltic pillows likely were comagmatic, albeit with different degrees of contamination. interestingly, an outcrop of nepheline syenite has been found in the southern slopes of flammefjeld (fig. 12), where the map shows quartz syenite of the snout series (author, unpublished observation). unfortunately this has never been followed up and its extent and relationships are unknown. the kangerdlugssuaq intrusion hosts many pegmatites, some merely coarse-grained equivalents of the main rock types, others agpaitic with minerals such as arfvedsonite, aegirine, astrophyllite, låvenite, catapleite and hjortdahlite. johnsen et al. (1998) described the mineral kentbrooksite from such a pegmatite, previously thought to be eudialite/eucolite (kempe & deer 1970). this is a mn rare-earth element – nb-f endmember of the eudialyte group. kupletskite, the mn analogue of astrophyllite, also occurs (christiansen et al. 1998), and the east greenland syenites seem to be unusually mn-rich (e.g. werner bjerge, where kupletskite has also been documented by brooks et al. 1982). another interesting mineral from these syenites is normandite (a mn-rich member of the cuspidine group), described by perchiazzi et al. (2000). a nepheline syenite pegmatite cutting nearby bagnæsset also contains astrophyllite (layne et al. 1982), an undetermined member of the eudialyte group, kupletskite and narsarsukite (c.k. brooks, unpublished data). astrophyllite also occurs on kræmer ø (see below). other syenites in the kangerlussuaq area (fig. 12) are much smaller in size and are all quartz-bearing. these include the kap boswell syenite and associated intrusions, the kap deichman syenite and associated bodies along hutchinson gletscher (fig. 13d; deer et al. 1984), the kræmer ø syenite (fig. 13e; brooks 1991a), the satellites of the kangerdlugssuaq intrusion (deer & kempe 1976), including the kærven syenites (holm 1991; holm & prægel 1988; nielsen 1989; holm et al. 1991) and bodies intruding the main kangerdlugssuaq intrusion. the bagnæsset quartz syenite may also be a satellite but was considered by deer & kempe (1976) to be earlier. many of these rocks have a peralkaline tendency (e.g. have late alkali pyriboles), others are thoroughly peralkaline with alkaline pyroxenes and amphiboles and minerals such as astrophyllite (layne et al. 1982) and aenigmatite (this most notably in the syenite of barberkniven in the kap boswell complex, fig. 12). characteristic accessory minerals of these sye nites are chevkinite and zircon. an almost aphyric, peralkaline vein cutting the kap boswell syenite, showing turbulence structures is shown in fig. 13f. a pegmatite belonging to the kræmer ø syenite contains large 4242 a d b f c g e trebjørnebjerget (foyaites) citadellet (pulaskites) 2 cm 2 cm 5 cm 10 cm fig. 13. syenites at kangerlussuaq. a: view from bagnæsset into amdrup fjord. the bagnæsset syenite, a satellite of kangerdlugssuaq intrusion, forms the foreground. trebjørnebjerget, c. 20 km distant, is at the centre of the kangerdlugssuaq alkaline intrusion (fig. 12). b: slab of kangerdlugssuaq foyaite. pinkish crystals: nepheline; off-white crystals: alkali feldspar; dark crystals: pyroxene and garnet. c: kangerdlugssuaq nordmarkite with 10–20% quartz (grey), sodic amphibole (black) and abundant miarolitic cavities. d: the kap deichman syenite intruding layered gabbros of the kap edvard holm complex (for location see fig. 9a). view from c. 1500 m altitude, looking west. e: contact of the kræmer ø syenite with archaean gneiss (left), showing part of the 10 m wide contact facies of breccia (right). f: almost aphyric, peralkaline vein cutting the kap boswell syenite, showing turbulence structures caused by the intrusion of two magmas with contrasting composition and viscosity. green colour caused by high aegirine content. g: breccia from the flammefjeld complex. large clast in the left seamed with molybdenite-filled fractures. 43 crystals of astrophyllite (layne et al. 1982), strongly sodic pyroxenes and amphiboles and a zone very rich in zircon and quartz. comendititic dykes which cut the skaergaard gabbros are thought to be late products of the kræmer ø syenite (brooks & rucklidge 1976). a massive biotite granite crops out on the south side of amdrup fjord. it has very widely spaced jointing and almost no dykes, so is clearly young. the flammefjeld complex will be referred to below. the felsic plutons generally contain few dykes and have widely spaced joint systems, which leads to massive walls and towers (fig. 13a). if the area were not so remote some of these rocks might produce handsome dimension stones: although sometimes blindingly white in outcrop (fig. 15a), the colour on freshly blasted surfaces is blue, resembling the well-known larvikite. occasionally, they may be crumbly, giving more rounded outcrops and areas of gravel. borgtinderne (fig. 2) is a major, largely syenitic body, as seen in fig. 6a invading the lower part of the cliffs in the main basalts (brown et al. 1978). along blosseville kyst, small felsic bodies are found in the ryberg fjord area and in the mountains to the east where syenite sills are quite common, although not studied in detail. they have yielded a u/pb zircon age of 50.2 ± 1.1 ma, a 40ar/39ar biotite age of 51.6 ± 0.04 ma, and a 40ar/39ar amphibole age of 54 ± 0.4 ma (p. japsen, personal communication 2011). the zircon and biotite ages are close to combined ages from a number of methods for the kangerdlugssuaq intrusion (see below), while the older amphibole age may be due to excess argon. in the i.c. jacobsen fjord area, syenite blocks found in the beach alluvium are thought to have been transported by glaciers and icebergs from the borgtinderne intrusion, which they closely resemble (brooks 1991b). small granitic and syenitic bodies also occur in the above-mentioned, largely mafic-ultramafic lilloise intrusion (brown 1973; chambers & brown 1995). a notable feature of almost all these intrusions is the presence of basaltic material. many contain angular xenoliths of basalt showing the former extent of lavas, even when no overlying basalt is present today. examples include the basaltic xenoliths in the kangerdlugssuaq intrusion, remarked on by all the authors cited above, and numerous large xenoliths in the kræmer ø syenite (brooks 1991a). others contain basaltic pillows and pillowed dykes, where ovoid basaltic bodies show crenulate, chilled margins to the enveloping syenite (fig. 17c). these are unequivocal evidence that mafic and felsic magmas were concurrently present. examples include well-developed pillows at astrofyllit bugt (nielsen & brooks 1991), studied in detail by riishuus et al. (2006), hybridising basaltic material at nuuk, kialeeq described by brooks (1977), shown in fig. 17c, and generally throughout the kialeeq district as described by brown & becker (1986), who mapped large areas as ‘net-veined complexes’. the latter name encompasses both pillowed bodies and others where inclusions are angular, although there is a complete transition between the two. also at borgtinderne (fig. 2), basaltic xenoliths and hybridising basaltic material were described by brown et al. (1978). a fine netveined/pillowed complex is found on the northern side of hutchinson gletscher (fig. 12) on the west side of kangerlussuaq, which may be part of the kap edvard holm complex (c.k. brooks, unpublished data). flammefjeld (geyti & thomassen 1984) is a unique locality in the kangerlussuaq area, although it closely resembles malmbjerg in the mesters vig area, to be described below. the top of the mountain is hydrothermally altered to brilliant red and yellow gossans (hence the name meaning flame mountain). although not exhaustively investigated, it consists of late-stage intrusive phases, aplite, quartz-feldspar porphyries, breccia pipes and dykes within the kangerdlugssuaq syenites. a breccia pipe contains numerous clasts, some of which display a stockwork of fractures containing fine-grained molybdenite (fig. 13g), a classical texture of porphyry molybdenum deposits. drilling, the only way to determine the depth and extent of this deposit, has not yet been carried out. a number of base and precious metal veins cut the surrounding rocks, notably the ‘yellow zone’ with tetrahedrite-bearing veins having elevated levels of copper, zinc, lead, silver and gold (thomassen & krebs 2001). interestingly, both the flammefjeld and malmbjerg occurrences show a close resemblance to the climax-type molybdenum deposits of the western united states, although the geological setting seems to be quite different. in addition to age data from the ryberg fjord area, reported above, several attempts to date the felsic intrusions have been made in the past (beckinsale et al. 1970; gleadow & brooks 1979; noble et al. 1988), but much of this earlier work is unreliable. as previously noted, k-ar dating was useful at the time, but it is not sufficiently precise to answer contemporary questions. tegner et al. (2008) reported high-precision 40ar/39ar results, showing that the oldest felsic rock is an alkali granite of kærven at 52.8 ± 1.3 ma, which cuts the 4444 24°w5 km a lp ef jo rd mestersvig a b pb-zn fle m ing fj ord kap simpson kap parry kap syenit werner bjerge traill ø malmbjerg oksehorn theresabjerg pictet bjerg 72°n kong oscar fjord st a u n in g a lp e r petersryggen centralen røde mur 72° n 24° w siriu s gletsc her schuchert g letscher taget hvide ryg ruinerne malmbjerg 24°w ar ctu ru s g let sc he r aldebaran gletscher drømmebugtenmeste rs vig mesoand palaeozoic sedimentary rocks basic intrusions acid volcanic rocks alkali syenite alkali granite nepheline syenite mesoand palaeozoic sedimentary rocks caledonian gneisses and granites palaeogene intrusions 25 km fig. 14. a: the mesters vig area showing three large subvolcanic complexes and five smaller ones forming a rough line oblique to the continental margin. these complexes contain large amounts of syenite and granite and werner bjerge also nepheline syenite. gabbro and diorite occur at theresabjerg, oksehorn and werner bjerge. b: map of the werner bjerge complex (from bearth 1959). location shown in fig. 2. 45 kærven gabbro (ohja 1966), itself dated at 55.1 ± 1.4 ma. the kangerdlugssuaq biotite granite, which crops out on the south side of amdrup fjord is considered, based on field relations, to be the youngest major intrusion in the area, has an age of 46.6 ± 0.9 ma (tegner et al. 2008). it was thought to be contemporaneous with the nearby porphyry molybdenum deposit at flammefjeld (geyti & thomassen 1984), but a re-os age on molybdenite of 39.6 ± 0.1 ma from flammefjeld was reported by brooks et al. (2004), showing it to be substantially younger. some alkaline basic dykes on the northern side of amdrup fjord, which contain xenoliths and megacrysts, are even younger: c. 34 ma (gleadow & brooks 1979). of particular interest is the summary of radiometric dates for the kangerdlugssuaq intrusion by tegner et al. (2008, fig. 3), which shows a constant age of c. 50 ma using a variety of methods (mineral and whole rock rb-sr, biotite k-ar, biotite 40ar/39ar, fission-track zircon and titanite) but a reduced age of c. 36 ma for fission-track apatite. this clearly shows that the intrusion rapidly cooled to a temperature estimated at c. 210°c and then more slowly to the closure temperature of apatite (c. 120°c). the study convincingly suggests that, in general, the ages obtained are ages of intrusion preserved by rapid cooling. the mesters vig area a further cluster of felsic rocks occurs north of scoresby sund (fig. 14a). no depth estimate exists for these intrusions, but it is generally agreed that they are very shallow, and they are regarded as subvolcanic. of these, the werner bjerge complex shown in fig. 14b is by far the largest and best studied (bearth 1959: general geology; brooks et al. 1982: mineral chemistry). here, a full spectrum of oversaturated, saturated and undersaturated types occurs, along with gabbros. the intrusions at kap parry and kap simpson (tyrrell 1932; schaub 1942) are caldera complexes with lavas (largely felsic porphyries interpreted as caldera fill, fig. 15b), but hydrothermal alteration is widespread (fig. 15c) and masks the rocks. some of the alteration haloes could mark blind ore deposits awaiting discovery. syenites and alkali granites and dyke swarms with a wide range of compositions (lamprophyres, trachytes and comendites, often with spectacular spherulitic textures) are also present. other smaller bodies of the area have been described by kapp (1960). their locations are shown in fig. 14a. of particular importance in this area is malmbjerg, originally known as ‘erzberg’ by the swiss and austrian geologists who discovered it. this is a small stock of perthitic granite with porphyritic aplite and granite intruding palaeozoic sedimentary rocks just west of the main werner bjerge complex that hosts the malmbjerg porphyry molybdenum deposit (figs 14a, b; kirchner 1964; harpøth et al. 1986; schønwandt 1988). the mountain is depicted in fig. 15d. with c. 150 mt of 0.23% mos2, malmbjerg is among the world’s largest molybdenum deposits and promises to be of exploitable value. however, access is difficult. the molybdenum deposit, a stockwork of quartz veins containing molybdenite, wolframite, topaz, fluorite, etc., forms an inverted bowl-like structure centred on the quartz porphyry. like the syenites at kangerlussuaq, those of werner bjerge have yielded a rich assortment of rare minerals (brooks et al. 1982): kupletskite, narsarsukite, pyrochlore, chevkinite, mn-pectolite, a li-bearing magnesioarfvedsonite, unindentified zirconosilicates possibly including hjortdahlite, as well as rosenbuschite. a new mineral, named kochite, which is a member of the rosenbuschite group, was described by christiansen et al. (2003). it occurs along with nepheline and alkali feldspar. k-ar and fission-track dating has been carried out of the mesters vig rocks over the years (beckinsale et al. 1970; rex et al. 1979; gleadow & brooks 1979). recent 40ar/39ar dating (c.k. brooks, c. tegner and r.a. duncan, unpublished data 2004) shows a spread of ages from 41.6 ± 0.5 ma for a syenite from kap parry to 26.9 ± 0.6 ma for the werner bjerge alkali granite and 25.5 ± 0.4 ma for the malmbjerg granite. these young ages reaching into the late oligocene were unexpected, but have been confirmed by a re-os age on molybdenite from malmbjerg of 25.8 ± 0.1 ma (brooks et al. 2004). this age distribution is obscure in the plate-tectonic context, although there seems to have been an acceleration in spreading rate at c. 30 ma (mosar et al. 2002b) and spreading on the ægir ridge eventually ceased at c. 25 ma. however, it is not obvious what relevance, if any, these correlations might have. geochronological data reported by price et al. (1992) on tholeiitic sills of the area indicate that they date back to c. 54 ma, giving a span of magmatic activity of nearly 30 ma, during which existing radiometric data do not suggest any clear gap, unless possibly in 4646 a c d 3 cmb fig. 15. a: foyaites (nepheline syenites) on the north side of aldebaran gletscher, werner bjerge. note sparse dykes, very little jointing, both indicative of young age, and mesozoic sedimentary rocks in mountain behind. the rocks have a very low content of ferromagnesian minerals and are very light coloured. b: typical feldspar-porphyritic trachyte such as makes up the caldera fill at kap simpson and kap parry (along with volcaniclastic deposits). c: extensive hydrothermal alteration in the caldera complex of kap simpson (from thomassen & nielsen 2006). d: malmbjerg seen from the schuchert gletscher with arcturus gletscher on the right. the granite cupola is deep red due to hydrothermal alteration that also extends out into the sedimentary host rocks (from thomassen & nielsen 2006). 47 the period 54–41 ma. it has been suggested by several authors that late activity in this area was linked to the splitting off of the jan mayen microcontinent from greenland by the northwards propagation of the kolbeinsey ridge (e.g. nunns 1983a, b), although, again, any link is very tenuous. the intrusions in this area apparently connect with an aeromagnetic anomaly offshore, which also appears on the conjugate vøring margin and was called the traill ø – vøring igneous complex by olesen et al. (2007). the mesters vig intrusions have also been linked by nielsen (1987) and larsen, h.c. (1988) to an abortive attempt to split off a continental fragment along a lineament roughly from inner kangerlussuaq to the werner bjerge area, as actually happened later, during the further rift jump with the jan mayen plateau. see also larsen, l.m. & watt 1985, although this work largely deals with the succession of basalt formations rather than the activity at mesters vig. as an alternative, could these intrusions follow the outer pseudo-fault (hey et al. 1989), which formed during the northwards propagation of the kolbeinsey ridge? the kialeeq (kialineq) – kap gustav holm area south of kangerlussuaq, felsic rocks are abundant in the kialeeq – kap gustav holm area (brown & becker 1986; brown et al. 1977; myers et al. 1993). the distribution of those in the kialeeq area (informally known as kialineq in the geological literature) is shown in fig. 16. much field work was carried out by the geological survey of greenland in the 1970s, but the results remain unpublished in the case of kialeeq. brooks (1977) described a striking example of hybridisation in the extensive net-veined unit. the intrusions at kialeeq include both syenites and granites, with extensive areas of acid–basic, net-veined complexes and the large imilik gabbro discussed previously. the area is poorly noe-nygaard intrusion 500 m 1041 m bjørn syenite kialineq diorite laubes gletscher syenite pueratse syenite nûk diorite ilitaaliip kangertiva langø nuuk ikâsangmît syenite tukingaleq store tindholm (takiseertivaq) lille tindholm (ittitalik) svineryggen kangikajik nuuluk laubes gletscher pilappik qajarsak granite uingaleq imilik kap hildebrand (imittip iliverta) a tter tiat kap s.m. jørgensen 600 m 1100 m 928 m 360 m 1942 m 66°45’n 67°n 34°w 10 km auluiartik granite 730 m kap warming imilik gabbro viewpoint fig. 17a d a n m a r k s t r æ d e gabbro/diorite archaean gneiss basaltic lavas syenite net-veined complex granite >5 0% d yk es < 5% d yk es 5– 50 % d yk es fig. 16. geological map of the kialeeq area (known as kialineq in the geological literature). intensity of dyking broadly indicated. geological localities shown in red, with corresponding geographical names in black (modern orthography). field of view of fig. 17a also shown. redrawn from bernstein & bird (2000) with additional observations by the author. location shown in fig. 2. 4848 a b c 49 known due to the extremely rugged nature of the landscape, with its extensive ice-cover, jagged mountains and coasts which almost everywhere rise steeply from the sea (figs 17a, b). radiometric ages have been reported by beckinsale et al. (1970), brown et al. (1977) and gleadow & brooks (1979), while a number of mineral and wholerock rb-sr isochrons, k-ar and 40ar/39ar dates have been determined by d.c. rex but remain unpublished. the two most recent are 40ar/39ar dates of 36.2 ± 0.6 ma from the nûk diorite (the rock studied by brooks 1977) and 37.2 ± 2.9 ma from the ikâsangmît syenite, both in the northern part of the area. previous dating suggests that intrusions in the southern part of the area are c. 2 ma younger (d.c. rex, personal communication 1998), but this suggested bimodality of ages remains to be tested. the reason for these young ages remains conjectural as the only major plate-tectonic event at this time was the cessation of spreading in the labrador sea (srivastava & tapscott 1986; roest & srivastava 1989) and a change from right-lateral shear to oblique divergence along the lena trough (engen et al. 2008). again, if there is any connection, it is not obvious. until the discovery of the sulugssut complex (see below under nephelinites), the kap gustav holm area was the most southerly known palaeogene intrusive centre in east greenland. only palaeogene sedimentary rocks and basalts were originally reported by wager (1934), but myers et al. (1993) described a number of generally circular intrusive bodies beginning with gabbro, which was tilted by the coastal flexure, followed by bodies of monzonite, syenite and granite, which are unaffected by the flexure. the granites gave a rb-sr isochron age of 53 ± 5 ma and a monzonite and the granite gave whole rock and biotite rb-sr ages of 50 ± 3 ma in both cases. k-ar ages of 51–49 ma were obtained for biotites and hornblende from these rocks. the complex is thus quite independent of the nearby kialeeq intrusions. syenite genesis in the hebrides and northern ireland on the other side of the rift zone, the predominant felsic rock type is granitic, found at most of the intrusive centres (e.g. emeleus 1955, walsh et al. 1979, meighan & gamble (1972). syenite only occurs in minor amounts (e.g. hole & morrison 1992). large amounts of felsic volcanic rocks have been reported from the volcanic successions of several other rifted volcanic margins (e.g. etendeka-paraná: peate, d.w. 1997; lebombo-nuanetsi: cox 1988; yemen-ethiopia: menzies et al. 2001). the felsic rocks in these cases are largely oversaturated rhyolites and/or granites, although some large igneous provinces, such as the siberian traps (zolotukin & al’mukhamedov 1988), columbia river (hooper 1997) and rajmahal (kent et al. 1997) seem to be relatively free of evolved rocks. phonolites do occur (e.g. marsh 2010, probably part of the etendeka province), but are not voluminous. perhaps the east greenland magmas with their largely syenitic compositions were less explosive than their rhyolitic counterparts in other provinces such that venting to the surface environment was minor. we do see, however, the classic distribution where volcanic rocks are largely basaltic and plutonic rocks are predominantly felsic, and there is a clear gap in sio2 contents (the ‘daly gap’). possibly felsic lavas were once more abundant in east greenland and have since been removed by erosion as was suggested by hunter & sparks (1987) to explain the low amounts of granophyre within the skaergaard intrusion. the conclusions of this paper were refuted by brooks & nielsen (1990), mcbirney & naslund (1990) and morse (1990), who among other arguments pointed out that no rhyolites are found in the lava pile. this makes the suggestion by hunter & sparks (1987) unlikely. erratic blocks of rhyolitic ignimbrite have been reported from the island of nordre aputiteeq (brooks & nielsen 1982a). their source is unknown, although they are likely related to nephelinitic parents such as the gardiner complex (the source of the numerous boulders of phonolitic tinguaites, probably deriving from dykes, found on raised beaches around the skaergaard facing page: fig. 17. geological features in the kialeeq area (known as kialineq in the geological literature). a: oblique air view of central kialineq: auluiartik granite at right centre (red arrow), northern part of imilik gabbro at bottom left. field of view shown in fig. 16. b: coast at the tindholm islands, showing the abrupt topography of the area and the difficulty of field work. c: basic pillows in syenite at nuuk where hybridisation in vertical pipes generates a diorite at the top of the picture (see brooks 1977), with hammer for scale. 5050 intrusion has likewise never been located; see brooks & rucklidge 1974). the origin of the east greenland felsic rocks remains conjectural. they were intruded over a much extended period (right down to c. 25 ma), which seems to rule out the possibility of derivation from the voluminous basaltic magmas whose extrusion was largely complete >20 ma before this. however, it must be borne in mind that basaltic magmas were available throughout the period of magmatism (albeit subsequently in small amounts), so this argument may be fallacious. two other possibilities have been entertained. firstly, that the syenites were produced by interaction of later basic magmas and their differentiates with country rock, as suggested above for the kangerdlugssuaq intrusion (riishuus et al. 2008), or secondly that underplated gabbroic material, produced during the main period of basaltic activity and imaged seismically (holbrook et al. 2001), became unstable, delaminated and was partially melted as it sank into the mantle to produce low-density syenitic magmas, which in turn invaded the upper crust. at present there are insufficient data, especially isotopic, to decide between these possibilities. as described above, recent information about the kangerdlugssuaq intrusion (riishuus et al. 2008) was interpreted as showing that it originated from a crustally contaminated, layered magma chamber at depth, followed by emplacement of the felsic magmas from the top of the chamber to their present position in the upper crust. in summary, problems related to the syenites include their large volumes, their young age (later than continental separation), the prolonged period of activity, the association of quartz-oversaturated with undersaturated types (perhaps now resolved by the work of brooks & gill 1985 and riishuus et al. 2008) and the tendency for many intrusions to be peralkaline, which seems at variance with the general aluminous nature of the continental crust and may be a serious problem for models with substantial crustal involvement in their genesis. furthermore, the genesis of two major porphyry molybdenum deposits remains obscure although they may be related to underlying amphibole-rich intrusions derived from mildly potassic, water-rich magmas which differentiated to silicic end products (the rhyolites and quartz porphyries at flammefjeld and mesters vig). nephelinites, carbonatites etc. due to their small volumes and often inaccessible locations such as in inland areas bordering the ice sheet, these rocks have hitherto been poorly documented, although a good deal of unpublished work exists, including a study by the author and stefan bernstein of the sulugssut complex. this probably explains why saunders et al. (1997) do not refer to them at all in an otherwise comprehensive review of the north atlantic igneous province. nevertheless, their presence is clearly significant, showing the products of low degrees of partial melting, perhaps under a thick lithospheric lid. the gardiner complex at the head of kangerlussuaq (figs 2, 18a) is the most impressive representative. it is the best described of this suite of rocks, having been documented in a number of papers by nielsen (1979, 1980, 1981, 1994), nielsen & buchardt (1985), nielsen & holm (1993) and nielsen et al. (1997). this ring complex is c. 5 km in diameter and consists of ultramafic cumulates cut by a later generation of melilitolites (fig. 18b) and carbonatites. it may be one of the most instructive occurrences of plutonic melilite-bearing rocks worldwide as exposure is excellent, although field work is offset by the frequent ferocious winds off the ice sheet. a study of the associated dyke rocks established the differentiation scheme for the parental melanephefacing page: fig. 18. nephelinites. a: the ring-shaped gardiner complex seen from the air, with the inner ring in the foreground (see fig. 12 for location: viewpoint approximately over batbjerg, looking west). b: layered melilite rock (uncomphagrite) from the gardiner ring dyke as described by nielsen (1980). dark layers are perovskite. yellowish colouring on the weathered surface may be due to cebollite, an alteration product of mellite. c: photomicrograph (crossed polarizers) of olivine nephelinite dyke rock from the batbjerg area, similar to material described by brooks & rucklidge (1974). field of view 4 mm. d: large, well-formed magnetite crystals, gardiner complex. e: bronze-coloured baryto-lamprophyllite, gardiner complex, as described by johnsen et al. (1994). f: photomicrographs of ijolite from sulugssut consisting largely of nepheline and augite rimmed by a na-rich variety of clinopyroxene. field of view 2 mm. left: plane polarised light. right: crossed polarisers. g: photomicrograph (plane polarised light) of phlogopite-rich nephelinite dyke-rock with a large mantled olivine, the sulugssut complex (brooks et al. 1989). field of view 4 mm. h: silico-carbonatite vein cutting gneiss at tuttilik. 51 e d b g fc a h 2 cm 2 cm 2 cm 5252 linite magma, in which early crystallisation of olivine and pyroxene led to the formation of the voluminous early cumulates (fig. 18c; nielsen 1994). further fractionation under conditions of both high and low fluid partial pressures and probably carbonatite immiscibility led to a variety of products, including melilitites, agpaitic phonolites and carbonatites. the intrusion apparently vented and dispersed phonolitic ashes over a very wide region (heister et al. 2001), as described above. melanephelinitic and related compositions, which are likely parent magmas of the gardiner complex, not only occur as dykes in the area, but also as lava flows at lindsey nunatak (nielsen 2001; peate, d.w. et al. 2003). loose blocks of nephelinite from raised beaches within the skaergaard area were described by brooks & rucklidge (1974) although it cannot be said whether these originated in flows or dykes. spectacular tinguaites with ≥1 cm large phenocrysts of nepheline besides alkali feldspar and katophorite, in a finegrained matrix rich in aegirine needles, occur in the same beach assemblage and probably also belong to this same genetic family. the gardiner complex is noted for its museumquality mineral specimens as described by johnsen et al. (1985). these include magnetite (fig. 18d), titanite, melanite (garnet), perovskite and apatite, all of which occur in large, well-formed crystals. gem quality titanite has also been recorded. detailed descriptions of titaniferous clinohumite (nielsen & johnsen 1978) and lamprophyllite/barytolamprophyllite, see fig. 18e, have also appeared (johnsen et al. 1994). 40ar/39ar dates from the gardiner complex indicate two periods of activity at 56.5 ± 0.3 ma and 54.7 ± 1.2 ma, probably corresponding to the field observations of early ultramafic cumulates and later ring dykes. the uncertainties arise because the ultramafic cumulates are not amenable to direct dating and only supposedly comagmatic rocks have been dated. these ages correspond to unpublished data by l.e. heister (2001) and rb–sr ages of 56.35 ± 0.24, 56.2 ± 2.4 and 56.14 ± 0.24 ma by waight et al. (2002), which were measured in situ using multicollector icp-ms and generating clinopyroxene-biotite, apatite-biotite and apatite-clinopyroxene-biotite isochrons, respectively. the sanidine-bearing tuff on the gronau nunataks c. 100 km to the west (see previous section on long-distance tephra falls), thought to be sourced at the gardiner complex, gives 55.0 ± 0.3 ma, the same, within error, as the second period of intrusion at gardiner (heister et al. 2001). ages of 55.0 ± 0.4 ma (sanidine) and 55.1 ± 0.5 ma (whole rock) were obtained on the same sample of this tuff by storey et al. (2007b, without acknowledging the earlier determination). within error this is the same age as petrographically similar tuffs from the north sea, denmark, the london basin and deep sea drilling project hole 550, as reported by heister et al. (2001). similar ages were also presented by westerhold et al. (2009). as previously discussed, this a very important marker horizon for the entire north-east atlantic. there is also a lu-hf age for the gardiner complex, using an apatite-whole rock isochron (barfod et al. 2003). however, the decay constant was poorly known at the time and the age was given as either 53.53 ± 0.53 ma or 51.84 ± 0.51 ma. the decay-constant problem now seems to be resolved (thrane et al. 2010) but the date still diverges from the others as it gives a value close to the younger estimate. a second nephelinitic intrusive centre, the sulugssut complex, lies some 200 km to the south (fig. 2). this is much less well known as only a short ‘discovery paper’ has appeared (brooks et al. 1989), but it contains ijolites (fig. 18f) and minor carbonatites. it is associated with intense dyke swarms in which tinguaites are prominent. they are aegirine-rich, both aphyric and with prominent nepheline, alkali feldspar and often titanite phenocrysts. nephelinites with mantled olivine also occur (fig. 18g). dykes probably related to this centre occur at tuttilik to the south (rucklidge et al. 1980). here are also found silico-carbonatite veins containing phlogopite and altered olivine (fig. 18h), which have been dated by storey et al. (2007b) at 58.3 ± 0.9 ma (two samples with identical results), considerably older than the gardiner complex. however, a gabbro from the sulugssut complex has given ages of 53.25 ± 0.27 ma (plateau age) and 53.30 ± 0.35 ma (l.e. heister, personal communication 2001). further work is needed to establish the relations between these nephelinitic/carbonatitic occurrences. a similar age was reported for a nephelinite flow at the base of the volcanic succession at hold with hope at 74°n by upton et al. (1989), although there is some doubt as to whether this might be a sill rather than a flow, as noted earlier (b. upton, personal communication 1998). within the plateau basalt area, a number of nephelinite vents were reported at sorte bræ by the danish lithosphere centre and they give a 40ar/39ar age of 50.2 ± 1.2 ma (tegner et al. 2008). nephelinite dykes have been reported from the interior of scoresby sund and, as already noted, as an isolated pyroclastic horizon 53 at the base of the rømer fjord formation (larsen, l.m. et al. 1989), which must have an age of c. 55.4 ma, judging by its position within the basalt stratigraphy. to the north at about 74° small flows and vents of highly alkaline lavas (melilitites, nephelinites and basanites) occur on the inland nunataks (fig. 2). these have been described by katz (1952a, b), brooks et al. (1979), bernstein et al. (2000) and harlou (2001) and have an age of c. 56 ma. they are in a pristine state of preservation, lack zeolites, appear never to have been buried and are found today in an environment where chemical weathering is close to zero as the temperature at this altitude (>1000 m) rarely rises above zero. they have exceptionally high contents of ti (up to 9%; similar rocks are only reported from bermuda. a description of the latter rocks by bernard gunn was never published but may be found at http://www.geokem.com/ oib-volcanic-atlantic.html#bermuda; (these same bermuda drill cores have also been studied in a master’s thesis by olsen 2005). these nunatak rocks were originally thought to be related to the jan mayen fracture zone as they fall near its extrapolation into the continent (brooks et al. 1979), a conclusion supported by torske & prestvik (1991), see below. nephelinite genesis several authors (marsh 1973; sykes 1978; garson & krs 1976; williams & williams 1977) have remarked on the localisation of kimberlites, melilitites, nephelinites, etc. on the landward extensions of transform faults in the adjacent ocean basins. indeed, in the north atlantic area torske & prestvik (1991) argued by that the vestbrona nephelinites on the norwegian shelf (møre platform, fig. 1) are found on an extension of the jan mayen fracture zone, following brooks et al. (1979) who suggested that the nephelinites of the nunatak zone of east greenland lay on the extension at the other end of this fracture zone. it is important to remember, however, that offshore structures are subject to interpretation. for instance, a series of apparent riftoblique transform faults in the faroe–shetland basin were, on closer inspection, shown to be caused by different processes including igneous intrusion (moy & imber 2009). larsen, h.c. (1990) pointed out that the east greenland shelf is segmented into distinct morphological regions characterised by differences in crustal type, contrasting basement tectonics, subsidence history and different timing of basin formation (fig. 19). these regions are separated by boundaries which appear to be some type of accommodation zones, which are, from south to north (fig. 19): the denmark strait escarpment zone at about 66°n (perhaps corresponding to the south iceland seismic zone at the present time), the kangerdlugssuaq escarpment zone at about 68°n (corresponding to the tjörnes fracture zone today and perhaps the ‘faroe transform fault’ of bott 1987), the scoresby sund fracture zone at about 70°n (spar fracture zone) and the jan mayen or kong oscar fjord fracture zone at about 73°n. on the basis of their limited distribution it seems very possible that the nephelinites, melilites and carbonatites of east greenland are related to these structures. sykes (1978) was of the opinion that oceanic fracture zones initiate at pre-existing weak zones in the continental lithosphere. this may determine the primary segmentation of the early rift, similar to that seen in east africa as described by rosendahl (1987; see also karson & brooks 1999, p. 334). it is clear that this suggestion concerning the distribution of nephelinites is wholly speculative and needs further work to be substantiated, and the absence of reports of nephelinites in the intervening areas does not necessarily mean they are not present, perhaps simply that they have not yet been found. a good test for the suggestion would be if nephelinites were found where they might be predicted, although no further discontinuities were mapped by larsen, h.c. (1990). if the average distance between the mapped zones is projected to the south it is to be expected that nephelinites might be found near ammassalik and further at umivik. however, none have yet been reported, with the exception of a nepheline syenite block at the old british arctic air route expedition’s base at nattivit, west of tasiilaq (formerly ammassalik) (unpublished observation by the author). nephelinites and melilitites form by small degrees of melting at relatively great depth. despite their small volume, these melts may have a significant role to play as they can potentially provide information about regions of the mantle that may not have been tapped by other magmas. furthermore, they are highly enriched in incompatible elements so that their signature would be very marked if mixed with other magmas, even in small fractions. they may in fact represent the enriched component of the icelandic mantle plume as suggested by bernstein et al. (2001). these authors suggested that 5454 a subducted oceanic slab might be involved, as required by those who deny mantle plumes (foulger et al. 2005). on present evidence, nephelinite magmatism seems to have been sporadic throughout the period of basaltic extrusion and no definite nephelinite event is apparent. dykes and sills cenozoic dykes and sills are ubiquitous in east greenland. precambrian dykes also occur cutting the gneisses, but these are easily distinguished by their metamorphic textures and commonly by the presence of garnet. in the caledonian areas north of scoresby sund, dykes of intermediate ages may also occur but are more difficult to recognise. the mica-rich monchiquite dykes of liverpool land referred to earlier (kofoed 1998) are an example. perplexing lamprophyre dykes described by larsen, p.-h. et al. (1990) are apparently truncated by upper permian sedimentary strata, which also contain clasts identified as coming from the dykes. nevertheless, the dyke rocks yield palaeogene ages. this phenomenon was explained by water-rich sedimentary rocks arresting the dyke intrusion. cenozoic dykes occur in many generations. some are intense, such as the coastal dyke swarm that cover almost 90% of the ground in extreme cases, while others are relatively sparse with spacings of many metres. some generations are regional (notably the coastal dyke swarm), whereas others are localised and associated with igneous centres, such as is probably the case with the ‘late dykes’ described by brooks & platt (1975). these dyke generations record episodes of crustal extension and changing magma compositions. in favourable circumstances, they can be placed into a relative chronological order by observing cross-cutting relationships both with each other and with their host rocks. sometimes the intensity of dyking and exposure can make the relationship between dykes very challenging to identify. it is not unusual to find dykes in areas of almost 100% exposure where precisely the intersection point is concealed by regolith. radiometric dating, likewise in favourable circumstances, can then place these generations into the absolute time scale. dykes therefore play a key role in documenting the tectonic and magmatic evolution. an exemplary study was that of nielsen (1978) of the dykes at kangerlussuaq. wager & deer (1938) made east greenland famous for its dyke swarm and flexure (fig. 20a, shown schea24 a21 a5 a7 a15 a29 a24 k o lb e in se y r id ge r e yk ja n es r id ge gfz ssfz kaez dsez tfz jmfz g r e e n l a n d 200 km sisz i c e l a n d sulugssut complex gardiner complex nunataks sorte bræ sydbræ gåseland aks e e bræ syydbdb and precambrian cratonic lithosphere caledonian folded lithosphere palaeozoic and mesozoic rifted lithosphere coastline shelf edge ocean-floor magnetic anomaly fracture/escarpment zone continent–ocean transition mid-ocean ridge nephelinite occurrence 69°n 32°w fig. 19. map of the east greenland margin, showing offshore features and, schematically, onshore geology. redrawn after larsen, h.c. (1990). gfz: greenland fracture zone. jmfz: jan mayen fracture zone. ssfz: scoresby sund fracture zone. kaez: kangerlussuaq escarpment zone. dsez: denmark strait escarpment zone. tfz: tjörnes fracture zone. sisz: south iceland seismic zone. the sisz and tfz are presumed to be the modern equivalents of the dsez and the kaez, respectively. larsen, h.c. (1990) stressed that major structural and stratigraphic changes take place across these zones as shown by aeromagnetic and reflection seismic mapping. positions of the main nephelinite occurrences in east greenland are shown with stars; a relationship to the zones is postulated. basalts are not shown. 55 a db c later generations early generation e fig. 20. dykes. a: coastal dyke swarm on the outer coast near miki fjord. note: apparent dips are distorted due to the orientation of the section. b: detail of the coastal dyke swarm at hængefjeldet as described by nielsen (1978). c: composite, c. 6 m wide dyke at bagnæsset. the outer parts are moderately potassic basanite and xenolith-free. the inner unit is crowded with a variety of kaersutite-bearing nodules and megacrysts likely representing an underlying cumulate body. d: dyke at wiedemann fjord containing occasional kaersutite megacrysts and numerous harzburgite nodules from the sublithospheric mantle. e: dolerite sill with palisade jointing on southern gauss halvø. dolerite sills are ubiquitous in the entire area of sedimentary basins. 5656 matically in fig. 21). similar dyke swarms are associated with flexures in the deccan and karroo provinces in india and southern africa, respectively. in this platetectonic era, it is clear that this structure represents the continental margin, where the dykes are equivalent to the sheeted dyke swarm present along mid-ocean ridges (e.g. karson et al. 1992) and that the flexure is a faulted structure bounding the newly formed rift. thus the term flexure, which suggests ductile deformation, is misleading. the dykes and sills are clearly a consequence of the extension and collapse of the continental margin (and are thus discussed again in several sections below). they also document the sometimes subtle changes in magma compositions with time. later studies of the dyke swarms in east greenland and the flexure associated with the coastal swarm have been made by nielsen (1975, 1978), myers (1980), nielsen & brooks (1981), karson & brooks (1999), callot et al. (2001, 2002), klausen & larsen (2002), klausen (2006) – all largely structural in approach; gill et al. (1988), hanghøj et al. (2003) – largely chemical; and by tegner et al. (1998b), lenoir et al. (2003), holm et al. (2006) – geochronological. some dykes in the wiedemann fjord area carry highly depleted lithospheric mantle nodules, which have been described by brooks & rucklidge (1973), bernstein et al. (1998a, 2006, 2007) and hanghøj et al. (2001). studies of specific groups of dykes were made by brooks & platt (1975) on the dyke swarm to the west of kangerlussuaq known as the ‘late dykes’ and the dykes at tuttilik (previously tugtilik) just south of 66°n (fig. 2; rucklidge et al. 1980). dykes that have suffered ca metasomatism were described by rose & bird (1994). 1.0 0 0.5 -0.5 post-flexure dykes syn-flexure dykes w (inland) e le va ti o n ( km ) faults 5 km e (danmark stræde) present topography b a present sea level pre-basaltic sediments palaeogene gabbro magmatic break-up unconformity ls ls ms m s us us site 989 site 917 site 916 site 915 site 990 site 914 1 km 0.5 1.0 precambrian gneiss palaeogene basalt pre-flexure dykespalaeogene syenite/granite eocene and post-eocene sediments tectonic break-up unconformity pre-basalt basement and pre-rift sediments 5 0 0 m tw o -w ay t ra ve l ti m e (s ) fig. 21. the coastal flexure of the east greenland continental margin. a: diagrammatic cross-section of the east greenland margin, showing flexured basalts and gabbros, numerous dyke generations and late felsic intrusions. b: interpreted seismic section across the margin showing some of the sites drilled on legs 152 and 163 of the ocean drilling program. redrawn from larsen, h.c. & duncan (1996). ls, ms and us refer to lower, middle and upper series basalts. 57 in a classic paper wager and deer (1938) described the coastal dyke swarm as following the coast for over 800 km from near scoresby sund to tasiilaq. larsen, h.c. (1978) discussed the offshore extensions of this swarm. wager & deer (1938) noted a constant relationship between three factors: the dip of the lavas, the density of the swarm and the dip of the dykes. thus, the densest part of the dyke swarm occurs where the lavas have the steepest dip, i.e. close to the coast. the dykes are nearly perpendicular to the lavas over the entire width of the swarm, suggesting that they have been tilted along with their hosts. wager & deer (1938) also noted that the coastal dyke swarm cuts the gabbros but not the syenites at kangerlussuaq or the syenites, the net-veined complex and the alkali granites at kialeeq, indicating its value as a time marker. the skaergaard, kap edvard holm and imilik gabbros are all tilted by the flexure and cut by the associated dykes. wager & deer (1938) suggested that the flexure arose by loading and sinking of the crust by 7 km of basalt and accentuated by melting and migration of sial to the west causing inland uplift, a model now obsolete. in his first paper, nielsen (1975) described the structure of the coastal belt, distinguishing early dykes, tilted seawards (i.e. landward-dipping) and later dykes with a more vertical attitude. these two major groups clearly bracketed the flexuring and break-up event. he further showed that the flexure involved considerable faulting, not a simple bending – the impression gained from earlier descriptions. later studies (e.g. karson & brooks 1999; klausen & larsen 2002) have refined this work. in a detailed study of the dyke swarms of the skaergaard intrusion area, nielsen (1978) recognised at least four generations, the first two being tholeiitic, the third alkaline and the final one containing both transitional and alkaline members. figure 20b shows an example of the complexity. he was able to tie their intrusion ages to the plutonic events and set up an igneous stratigraphy spanning from c. 58 to c. 36 ma, using the then sparse radiometric data. on the basis of chemical similarities gill et al. (1988) claimed that nielsen’s earliest dykes were contemporaneous with the lower basalts, which we now know to be c. 60 ma old. using the same criteria they claimed that other dykes were largely feeders to the overlying plateau basalts. a much more detailed study of the elemental and isotopic compositions of the dykes was made by hanghøj et al. (2003), although correlation with nielsen’s generations is not always apparent. karson & brooks (1999) stressed the segmentation of the coastal region in the north–south direction, showing how the earliest phase of flexuring entailed considerable penetrative deformation including the formation of mylonites and pseudotachylites, dated by karson et al. (1998) to 62.9 ± 4.5 ma using laser heating techniques. these pseudotachylites are clearly evidence that at times the margin was magmastarved and the extension was largely tectonic. later on, much more voluminous magmas were produced and extension was accomplished entirely by dyke injection as at normal mid-ocean ridges, and the tectonic deformation testified by the pseudotachylites ceased. both tectonic and magmatic styles of extension are known from mid-ocean ridges. klausen & larsen (2002) used statistical methods to reveal the structure of the dykes and the flexure, by showing the relation between dyke swarms and igneous centres, as in iceland. lenoir et al. (2003) dated tilted dykes at c. 55–54 ma and vertical dykes at c. 51 ma at kap wandel (66°20´n) and kap gustav holm (66°40´n) and were able to estimate a minimum strain rate, which suggested a high thermal gradient. magnetic anisotropy studies (callot et al. 2002) indicate that tholeiitic dolerites were injected by horizontal flow, as would be expected if they were fed from the major centres, as suggested by klausen & larsen (2002) and also similar to the horizontal flow documented for krafla in iceland (björnsson 1985). probably the best site to study the present splitting of continental crust to form a new ocean is the woodlark basin of papua new guinea, which has been investigated by a variety of methods including drilling by the ocean drilling program (speckbacher et al. 2011 and references therein). here, extension takes place along low-angle detachment faults with a kilometre-scale displacement prior to the inception of volcanism and the creation of new ocean-floor. these faults are associated with extensive cataclastites and mylonites and lubricated by talc, comparable to the deformed gneisses and hyaloclastites described from east greenland by karson & brooks (1999). volcanism follows closely after this tectonic extension with the formation of new ocean floor. it may well be that the woodlark area provides a useful modern analogue to the events that took place during the earliest phase of break-up in east greenland. there is a divergence of opinion regarding the precise mechanism for the emplacement of the east greenland coastal dyke swarm. wager & deer (1938) and wager (1947) envisaged the dykes as being intruded in a fan-shaped array in association with the flexuring; while nielsen (1975) considered that the dykes were in5858 truded vertically in the basalt pile and then tilted by the flexuring. faller (1975) and faller & soper (1979) carried out palaeomagnetic measurements in the miki fjord – irminger fjord area and found that their results precluded later tilting of the dykes and that they were intruded in their present orientation. this question remains unsolved, mainly because of the high levels of alteration in the area sampled by faller & soper (1979) and also because it was not appreciated at the time that the basalts in this area (lower basalts) were appreciably older than the dykes cutting them. it is thus suspicious that the dykes and the flows have the same palaeomagnetic directions and that these are also the same as the youngest flows in the main basalts as reported by tarling (1967) and hailwood (1977). a suite of dykes trending roughly n–s in the middle reaches of the western side of kangerlussuaq (fig. 20c, known as the ‘late dyke swarm’ as it cuts everything else in the area) was studied by brooks & platt (1975). these dykes make up a differentiated mildly potassic suite and are remarkable for their kaersutite-bearing xenoliths, often in large size and quantities. brooks & platt (1975) concluded that they are evidence for a major layered kaersutite gabbro intrusion under this area. some dykes here are as young as 35 ma (gleadow & brooks 1979). there are outcrops of a kaersutitebearing dioritic rock, often rich in large crystals and aggregates of pyrite, on the ridge between sdr. syenit gletscher and the crags on the south side of ndr. syenit gletscher, which may be part of this otherwise buried body. the possibility that the quartz porphyries of flammefjeld are related to this blind intrusion should be examined. brooks & platt (1975) showed that such oversaturated, silicic magmas could be generated by fractionation of amphiboles, which are relatively silicapoor. dykes cutting the skaergaard intrusion were described by vincent (1953). they range from dolerites to camptonites (hornblende lamprophyres), which appear very prominent due to light weathering colours. vincent (1953), whose samples were collected within a restricted area in the south-western part of the intrusion, was unaware of the comendites which cut the northern part of the intrusion (brooks & rucklidge 1976). these comendite dykes are c. 1 m thick and are of two types. one contains phenocrysts of quartz, alkali feldspar and amphibole in an aegirine-rich groundmass. the bulk of the dyke is not peralkaline and buff-coloured in outcrop, whereas the margins are peralkaline and bluish in outcrop. the more slowly cooled, buff material has clearly lost alkalies, and in a thin section, the arfvedsonite needles of the margin are seen to be oxidised. the other type of dyke is aphyric. it is bluish over the entire width, due to high levels of acicular alkali amphibole in the groundmass, and is markedly spherulitic. in the porphyritic dykes it is thought that crystallisation has been caused by seeding on the contained crystals, whereas this has not taken place in the aphyric varieties, which were quenched below the glass transition and devitrified over time. where such dykes cut the gabbros of the skaergaard intrusion the adjacent gabbros have been strongly bleached over a few centimetres. these dykes are thought to be related to the kræmer ø syenite. similar dykes occur throughout east greenland, e.g. associated with the kap simpson/ kap parry intrusions to the north (schaub 1942). an orbicular-structured comendite, belonging to the aphyric group, was described by mcbirney et al. (1990) as an example of self-organisation. they referred this dyke to the skaergaard transgressive granophyres although peralkaline dykes of this type are quite common in the northern part of the intrusion and its host rocks and appear to radiate from the kræmer ø syenite. one prominent very early dyke cutting the skaergaard intrusion contains gabbroic xenoliths thought to represent deeper, unexposed levels of the intrusion (irvine et al. 1998; jakobsen et al. 2010). compositionally, it resembles the postulated skaergaard initial liquid. of major interest are dykes at wiedemann fjord on blosseville kyst. two of these dykes, occurring in the north-south fractured zone of basalts mentioned previously, have been dated at 41.5 ± 0.9 and 36.6 ± 0.5 ma (40ar/39ar, tegner et al. 2008). the dykes are camptonites and perhaps also monchiquites, and some contain a suite of amphibole/pyroxene megacrysts and nodules of strongly depleted harzburgite (brooks & rucklidge 1973; fig. 20d), which were shown by bernstein et al. (1998a, 2006, 2007) to represent residues from partial melting in which komatiitic magmas were produced. the nodules give a unique indication of the nature of the subcontinental lithosphere in this region, and are unique in the north atlantic igneous province with the exception of similar material collected from ubekendt ejland, west greenland, where they occur in a basanitic lava flow younger than 52.5 ma (bernstein & brooks 1998). using osmium isotope model ages hanghøj et al. (2001) were able to confirm that the melting event generating komatiites took place in the archaean. similar material has been described from kimberlites and ultramafic lamprophyres in west 59 greenland and canada (bizzarro & stevenson 2003; wittig et al. 2008). other dykes to be found are composite dykes with mafic margins and felsic interiors, common in kialeeq, and felsic dykes with mafic pillows, observed at kap boswell and kruuse fjord. many of the syenites contain synmagmatic dykes, i.e. dykes broken up to variable extents as they were intruded into the still unconsolidated host. sills and sill complexes are common in the areas of sedimentary rocks. the two most prominent host areas are the jameson land basin at scoresby sund and the sorgenfri gletscher area of the southern blosseville kyst, although sills occur throughout the sedimentary basins of north-east greenland (fig. 20e). extensive sills also occur on traill ø accompanied by a small outcrop of basalt (hald 1996). the jameson land suite of sills and dykes, trending ese, has been described by larsen, h.c. & marcussen (1992) and hald & tegner (2000). these sills are mainly evolved plagioclaseaugite-olivine-phyric tholeiites, and are cut by alkaline dolerite sills and dykes in jameson land. surprisingly, this sill complex post-dates the plateau basalts by 2–5 ma and the initiation of sea-floor spreading by 1–2 ma. the sorgenfri gletscher sill complex (wager 1947), which intrudes the sedimentary rocks of the kangerlussuaq basin, has been studied in a phd thesis (gisselø 2001). individual sills may be several hundred metres thick. they are of tholeiitic composition and can be compositionally matched in the plateau basalts, and 40ar/39ar dating shows that they are contemporaneous with the basalts (tegner et al. 1998b). there are also a number of syenitic sills in this area, which corresponds to a dome-shaped area of high palaeo-temperatures revealed by fission-track studies (hansen & brooks 2002, see fig. 22). a specimen containing aegirine and arfvedsonite from a nunatak that has recently emerged from the ice at the snout of sorgenfri gletscher (fig. 2) was described by brooks (1991a). the age of these sills has already been given in the syenite section above. similar sill complexes have been noted by several authors (e.g. hansen 2006) from the conjugate margin in the rockall, faroe–shetland and møre basins. hansen (2006) showed that their intrusion took place in several events. this led to venting through the overlying sedimentary rocks, which was proposed by svensen (2004, 2010) to have led to massive methane additions to the atmosphere and caused the paleocene–eocene thermal maximum (see also the section, time frame for magmatism and margin deformation). cutting the skaergaard intrusion and the coastal area to its south are some large sills, which have been described by hughes (1956), douglas (1964) and naslund (1989). the basistoppen sill which cuts the skaergaard intrusion (originally interpreted as a raft of basalt by wager & deer 1939), is 660 m thick and displays one of the most complete differentiation sequences known, comparable to those of the skaergaard and bushveld intrusions. this may be because of intrusion into a hot host with consequent slow cooling. its margins also show rheomorphic effects. recent, highly precise 238u/206pb zircon dates, using the id-tims method, show that the sill was intruded 125 ± 85 ka before the closing temperature for zircon of the skaergaard intrusion’s sandwich horizon: the uppermost unit of the layered series (wotzlaw et al. 2011). large dykes up to 1000 m thick also occur in the skaergaard area. they are termed macrodykes and have been described by bird et al. (1985), white et al. (1989), blichert-toft et al. (1992), geist & white (1994) and momme & wilson (2002). they are thought to be comagmatic with the skaergaard intrusion and show layering, abundant xenoliths and autoliths, rheomorphic effects (e.g. naslund 1986) and development of granophyres and pegmatites. the kræmer ø macrodyke was interpreted by momme & wilson (2002) as a feeder to the overlying basalts. the miki fjord and vandfaldsdalen macrodykes have been the subject of studies of reaction and hybridisation with the country rocks (rosing et al. 1989; blichert-toft et al. 1992; waight & lesher 2010), which have implications on the mechanism of contamination seen in some of the lavas, particularly the lower basalts. these macrodykes also contain interesting concentrations of noble metals (brooks et al. 1987) and have recently become the target of mineral exploration by platina resources limited. a suite of dykes described by brooks & nielsen (1982b) is also comagmatic with the skaergaard intrusion. these dykes contain granophyric veins with a composition similar to plagiogranite of ophiolites (c.k. brooks, unpublished data) and were used to illuminate the early phases of compositional development of the skaergaard intrusion’s magma. two other types of dykes should also be noted, namely a group bearing gneiss xenoliths discussed by bird et al. (1985) and those bearing green, often spectacular rosettes of prehnite up to 3 cm across (bird et al. 1985; rose & bird 1994; see also below). the gneiss-bearing dykes were thought to follow the e –w topographic low 6060 along forbindelsesgletscher in the skaergaard intrusion, inner miki fjord and the glacier between miki fjord and i.c. jacobsen fjord. the metasomatic processes taking place in the prehnite-bearing dykes have been examined in detail by bird et al. (1985) and rose & bird (1994). the coast-parallel dyke swarm and the closely associated coastal flexure are critical to an understanding of precise break-up mechanisms, as addressed in many of the references cited above. the seaward increase in the intensity of dykes and their increasing dip are a striking example of extension and continental margin collapse during continental rifting. the lack of a similar dyke swarm on the conjugate margin suggests that break-up was asymmetric as was suggested by wernicke & tilke (1989) to be the case for the eastern north american margin. however, as the conjugate margin is submarine, exploration is incomplete and such structures are difficult to identify by geophysical methods. the presence of pseudotachylites, sometimes in abundance, indicates that at times extension was largely tectonic, rather than magmatic. observations suggest that the pseudotachylites were formed at an early stage of rifting before large volumes of magma were available, as seen in the woodlark basin referred to above, but the pseudotachylites and their relations deserve a more detailed study. the woodlark rift is strongly asymmetric. recently the east greenland margin and its conjugate hatton bank margin (west side of rockall plateau, fig. 1) were investigated by white & smith (2009), who confirmed the asymmetry. they suggested that this may have started at an early stage before copious amounts of magma were available, and that the structure was at this time similar to non-volcanic margins, which are conspicuously asymmetric. whether the rift is symmetric or asymmetric has important implications for the break-up and subsequent spreading. finally, the segmentation of the margin as described by myers (1980), karson & brooks (1999) and klausen & larsen (2002) and shown in fig. 19 gives a powerful insight into the nature of continental break-up and invites comparison with segmentation of the mid-ocean ridges (macdonald et al. 1991) and continental rifts (e.g. rosendahl 1987; beutel et al. 2010). to summarise, the dykes of east greenland are important because they provide time markers, indications of the directions and times of extension as well as the changing compositions of the magmas. furthermore, they sample unexposed material from the subcontinental lithosphere, from concealed intrusions and from unexposed crust (many dykes in areas of basalt contain gneissic inclusions). much of the history of the margin can thus be interpreted from the dyke swarms. hydrothermal activity patches of intense hydrothermal alteration are to be found in association with the subvolcanic complexes of the mesters vig area (figs 14a, 15) and in particular at the malmbjerg molybdenum deposit. likewise, the flammefjeld molybdenum deposit at kangerlussuaq owes its name to the bright red and yellow colouring of the mountain top due to hydrothermal alteration. this is probably to be expected as intrusions north of scoresby sund were emplaced into sediments and would readily have set up hydrothermal systems, whereas this does not happen when the country rocks are gneiss of low permeability, as documented for the skaergaard intrusion (norton & taylor 1979; taylor & forester 1979). flammefjeld thus seems to be an exception. basalts also are reasonably permeable on a macroscale and may be cut by hydrothermal veins. bright red calcitic veins, up to several metres wide, cut the basalts sporadically throughout the area (e.g. on the low rocky outcrops beyond the head of wiedemann fjord, see fig. 6h). the ability of basalts to support hydrothermal systems has been documented by manning & bird (1991, 1995) and manning et al. (1993). high permeability is maintained via the slaggy flow tops and bottoms, until it is eventually reduced by the deposition of zeolites and other minerals. thus, hydrothermal circulation in the basaltic pile has led to the zeolite zonations described by neuhoff et al. (1997) referred to above. copper-prehnite-calcite assemblages (bird et al. 1985) belong to greatest depths exposed and native copper may occur in vesicles in the basalts, associated with calcite, prehnite, etc. such copper seems to have been utilised by the inuit, as artefacts from a recently found house site at søkongen ø just south of nansen fjord include small copper rivets fastening baleen strips that bind coopered vessels. these abundant artefacts were found by the hunter asser johansen and are now in the national museum in nuuk. to the author’s knowledge this indigenous use of native copper in greenland has not been commented in the literature, although the use of meteoric and telluric iron is well known (buchwald & mosdal 1985). 61 following the pioneering work of taylor & forester (1979) on the skaergaard intrusion, numerous careful and detailed studies have illuminated the circulation of hydrothermal waters in the intrusions and their country rocks by mapping fracture systems and studying the mineral assemblages produced: e.g. bird et al. (1985, 1986), manning & bird (1986, 1991, 1995), rose & bird (1987, 1994), fehlhaber & bird (1991), manning et al. (1993), nevle et al. (1994) and brandriss & bird 1999. the last of these publications studied some remarkable dolerite dykes, which had been subjected to ca-metasomatism and contain gem-quality, green prehnite rosettes in their central parts, as referred to in the last section. showings of hydrothermal sulphides occur at a large number of places north of scoresby sund and have been described in a major monograph by harpøth et al. (1986). many of these are associated with the palaeogene intrusions, such as werner bjerge, or are thought to be related to other palaeogene activity. others are of uncertain age, such as stratabound cu-pb-zn deposits on wegener halvø. similar sulphide occurrences are associated with the intrusions in the kangerlussuaq area. in particular, base and precious metal veins with galena, sphalerite, chalcopyrite and tetrahedrite, related to the flammefjeld pipe, were described by thomassen & krebs (2001). as the host rocks at flammefjeld are also likely to be of low permeability, being syenites with large fracture spacing, the hydrothermal alteration and mineralisation are probably owed to an unusually water-rich magma. to recall a point made previously in the section on felsic intrusions, flammefjeld is closely associated with an intrusive body that is rich in amphibole (kaersutite, as documented by brooks & platt 1975) and therefore water-rich. the quartz porphyry magma is also likely to have been volatile-saturated although advanced alteration does not allow its primary mineralogy to be discerned, beyond quartz phenocrysts. such volatile-rich magmas would have the ability to fracture the country rocks and set up their own hydrothermal systems. the youngest representatives of the ‘late dykes’ (see previous section) are quartz normative (brooks & platt 1975) and may be equivalents to the quartz porphyries of the flammefjeld system (as noted previously). an oxygen isotopic investigation of flammefjeld is overdue. similar dykes are present at malmbjerg (author’s observation) and this deposit may have a similar origin. in an early attempt to characterise these deposits, lead isotopes were investigated by coomer et al. (1974) with a view to aid the exploration efforts of the nordisk mineselskab a/s. more recently, jensen (1998) continued this work, concluding that the deposits were mixtures of metals from the nearby country rocks. thus, at kangerlussuaq, palaeogene lead of distinct north atlantic character has mixed with primitive lead derived from the surrounding gneisses. north of scoresby sund, however, the lead isotope composition of igneous-related deposits shows that they were derived by mixture with proterozoic and caledonian sources. while the stratabound deposits of wegener halvø are thought to be of palaeogene age, the lead has been derived from the thick, palaeozoic-mesozoic sedimentary rocks of the jameson land basin. jensen’s work also included an extended discussion of the lead systematics of the igneous rocks themselves. it should be noted that there is little evidence that the skaergaard intrusion’s precious metal deposit (or those in nearby gabbros) is hydrothermal. although its precise genesis is unknown, most workers favour a purely orthomagmatic origin (bird et al. 1991; andersen et al. 1997; nielsen et al. 2005). geomorphology the first impression one gets of east greenland is its mountains, which stretch along almost the whole length of the coastline and attain heights of almost 4 km. what is the origin of this mountainous terrain and how long has it been elevated? the author (brooks 1979, 1985a) studied the topography of the kangerlussuaq area in some detail and offered explanations for the observations. although geomorphology is generally given scant attention by hard-rock geologists, it can make a considerable contribution; not only regarding the contrasting erosion of different rock types and the development of landscapes, but also to changing elevations, which may reveal more fundamental processes taking place in the crust or mantle beneath (e.g. clift et al. 1998). this idea has recently gained much more appreciation and, for example, previous uplift by the supposed icelandic mantle plume has been documented in buried landscapes west of the shetland islands (lovell 2010; hartley et al. 2011). the striking topography of east greenland is clearly related to the large-scale effects of continental break-up and is probably affected by the impingement of the postulated plume at the base of 6262 a c d b gardinergardiner >50 ma 40–50 ma 30–40 ma no major intrusions exposed 0123 4 5 6 borgtinderne 2.1 2.6 2.5 2.5 2.5 20–30 ma <20 ma blosseville kyst kangerlussuaq 30–40 ma ? 3 4 contours on dome contours on apatite fission-track ages 2 2 50 km 2.5 preserved uplifted plateaux all heights in km e fig. 22. geomorphological features. a: the lemon bjerge seen from the foot of frederiksberg gletscher. the c. 2700 m high summits of these mountains reach almost to the original gneiss-basalt unconformity and trace the outline of the kangerlussuaq dome. b: coast near miki fjord. the basalts along the southern part of the blosseville kyst have been deeply buried and intensely intruded by dykes and have lost the typical trap topography prevalent elsewhere and seen in fig. 6a. c: exhumed pre-basaltic peneplain just north-east of the head of kangerlussuaq and north of lemon bjerge, representing the core of the kangerdlugssuaq dome. nordfjord gletscher in the foreground, watkins bjerge in the distance. d: escarpment of the watkins bjerge (c. 3700 m), representing the eroded edge of the updomed basaltic cover. e: the kangerlussuaq dome (for method of reconstruction see brooks 1979). numbers on black contours show the estimated height of the original basalt surface in kilometres. numbers in inland areas give the height of preserved, uplifted peneplains. red curves show the distribution of apatite fission track ages (hansen & brooks 2002) revealing a domed area of enhanced heat-flow, somewhat offset from the physical domal structure. 63 the lithosphere. there are thus two superimposed effects which must be considered: one causing horizontal movement (continental break-up), the other largely responsible for the vertical movement (epeirogenesis). it should be recalled that iceland is not only regarded as a classical hotspot or plume site, but is also centred on a spreading ridge that is anomalously elevated (vogt 1974). a better understanding of the origins of the east greenland topography is expected to provide valuable clues relating to the formation of passive margins. a major stumbling block is, however, that geomorphological processes, particularly uplift events and the formation of peneplanes, have been difficult to date. thus, if the crustal uplift was at a markedly different time from the plume activity, we would need to reject the plume hypothesis for uplift. the raised surfaces could not generally be dated prior to the advent of fission-track dating. increased elevations can often be dated by the timing of sediments shed into adjacent basins, but these have not been explored in east greenland. whilst little work has been published on the geomorphology of the area north of scoresby sund since the pioneering work of ahlmann (1941), the geomorphology of the kangerlussuaq area was addressed by the author (brooks 1979, 1985a), who found that this area is characterised by three distinct domains (fig. 22): the coastal flexure, the inland plateaux (increasingly dissected and descending towards the coast) and a major domal structure centred on the gneissic mountains, lemon bjerge, which culminates in domkirkebjerget (fig. 12) inland from the skaergaard intrusion. to these must be added the additional features in the basalt area nearer to scoresby sund, referred to above (pedersen et al. 1997), namely a local point-source uplift possibly indicating an underlying pluton, a half-graben with the igtertivâ lavas and a structurally complex zone paralleling the coast. traditionally, a great drawback to geomorphological studies has been the difficulty of dating typical landscape features, as noted above. the introduction of fission track and (u+th)/he dating in the last c. 40 years has given rise to an increasing number of studies, although interpretation is often not straightforward. it is important to note that these methods are not generally aimed at finding the formation age of a rock, but are designed to reveal its thermal history, which, in turn, tells us something of the burial and uplift history of the sample. fission-track ages may nevertheless reflect the crystallisation age of a rock if cooling was rapid (e.g. for a magmatic rock, its crystallisation age), albeit with lower precision than with other methods. ages are derived from the density of fission tracks combined with the concentration of u, but as the tracks spontaneously shorten over time, depending on the ambient temperature, track lengths must be taken into account to understand the thermal history, and this usually involves numerical modelling. the many pitfalls inherent in the interpretation of such data have been described by redfield (2010) using publications from west greenland as an example. this work is very relevant to the discussion of the east greenland uplift history as the two areas show many features in common and detailed work on east greenland is not yet available. redfield (2010) denied that the case for neogene uplift can be sustained by the available data, but redfield’s criticisms were robustly refuted by green et al. (2011). they emphasise that their results show the following events: paleocene subsidence, eocene–oligocene uplift (36–30 ma), further uplift at c. 10 ma and final uplift in the pliocene (7–2 ma), and that these conclusions are backed up by consistent geological, vitrinite reflectance and apatite fission-track studies. green et al. (2011) made the important observations (also made previously by brooks 1979) that the later events are not directly related to continental break-up and that mechanisms to explain them are lacking. an early application of fission-track studies in east greenland was that of brooks & gleadow (1977) to derive an age for the skaergaard intrusion, obtaining 54.6 ± 1.7 ma for zircon. disregarding the large error, the result is essentially identical to more precise ages obtained with u-pb methods by hirschmann et al. (1997), hamilton & brooks (2004) and wotzlaw et al. (2011). apatite u-pb data from the same sample as the zircon gave 38.1 ± 3.1 ma. this indicates that there was protracted cooling, reflecting uplift and erosion of the overlying basalts over c. 20 ma. a later comprehensive study by gleadow & brooks (1979) on a regional basis established that this late cooling took place within the domal structure described by brooks (1979; see fig. 22) and probably represents a combination of high heat flow and subsequent unroofing. nevle et al. (1994) used these observations to postulate palaeogene climate change in the area, explaining oxygen isotope values in hydrothermal fluids as arising from drainage of waters with falling δd from increasingly higher altitudes as the dome rose from c. 55 to c. 50 ma. gleadow & brooks (1979) interpreted their results in the sense that the regional plateau uplift had oc6464 curred at c. 35 ma, coincident with the intrusion ages at kialeeq, but whether this correlation is meaningful remains obscure as there are very few intrusions of this age at kangerlussuaq. gleadow & brooks (1979) also extended their study to the caledonian region to the north showing protracted cooling due to slow denudation of the mountain range. they also dated some of the mesters vig intrusions, finding that they were emplaced over a large time span, extending to unexpectedly young ages (<30 ma) and cooled rapidly, as would be expected for such shallow-level, indeed subvolcanic, complexes. ages of down to 25 ma were later confirmed by more precise methods (e.g. re-os by brooks et al. 2004) there have subsequently been a number of fissiontrack reports: hansen (1992, 1996, 2000), clift et al. (1998), thomson et al. (1999), mathiesen et al. (2000), hansen et al. (2001), hansen & brooks (2002) and hansen & reiners (2006). the last of these studies includes (u+th)/he data which showed discrepancies relative to apatite fission track results. the geomorphology gives irrefutable evidence that post break-up uplift has occurred, but as to the precise timing the controversy between redfield (2010) and green et al. (2011) in west greenland is instructive. it is now generally agreed that significant uplift occurred at least as late as the miocene in broad areas, which makes it unlikely that uplift was caused by underplating as the supposed plume was remote at this late stage (clift et al. 1998). hansen & brooks (2002) found an interesting pattern of ages centred inland from nansen fjord and, rather surprisingly, somewhat offset from the dome reconstructed by brooks (1979) based on geomorphological considerations (fig. 22). this is an area where cooling has taken place very late but it does not obviously correlate with any other geological event in the area – the sorgenfri sill complex is much older (tegner et al. 1998b). perhaps the syenite sills in this area are relevant as they may indicate an underlying, blind pluton, but their dates (quoted in the section on felsic intrusions above) show that these are also too old. fission-track dating has shown that inland areas have been quiescent since the mesozoic and no trace was found of a plume passage heading inland from the head of kangerlussuaq fjord (hansen 1996) as might have been predicted from existing models. north of scoresby sund, the country was divided into crustal blocks which behaved independently of one another, giving the present relief dominated by halfgrabens. much of the development of this area can be deduced from the mesozoic stratigraphy (e.g. surlyk 1990). net uplift since the carboniferous has been c. 9 km for the stauning alper, also a site of domal uplift (ahlmann 1941), and, in contrast to the situation farther south, cooling has been rapid, but complicated, from c. 55 ma ago (thomson et al. 1999). it seems that fission-track and (u+th)/he analyses will lead to valuable insights into the development of the landscape, but at present such studies are too sporadic and disconnected to develop a comprehensive picture. moreover, there are still disputes regarding interpretation. perhaps the most significant result to date is that the general uplift of east greenland and probably other areas of the north atlantic igneous province, which is one of its distinguishing features and part of vogt’s (1974) ‘icelandic phenomenon’, is the recognition that the uplift has been relatively recent. however, this conclusion is by no means certain. future aims will be to better constrain the timing of the uplift and, if possible, identify its mechanism. such a study is being undertaken at present (p. japsen, personal communication 2010). structure of the east greenland margin and its setting within the north atlantic region in this section the general structure of the east greenland margin and how it relates to ocean floor initiation will be summarised and discussed. in many ways it will be a summing up of what has already been said, but it is also an attempt to put east greenland within a broader setting. the nature of volcanic margins worldwide, with some examples from east greenland, has been discussed by geoffroy (2005); this paper gives an overview of contemporary thinking about such margins. the coastal flexure and its dykes in summary, the structure of the east greenland margin is dominated by the coastal dyke swarm and associated flexure first described by wager & deer in 1938 and later discussed by nielsen (1975, 1978), myers (1980), nielsen & brooks (1981), karson & brooks (1999), klausen & larsen (2002) and klausen (2006) 65 – see also rucklidge et al. (1980) for a description of the dyke swarms at tuttilik in the southern part of the area. cataclastic deformation and generation of pseudotachylite (karson et al. 1998) represent the early stage of crustal extension. the flexure and its dyke swarm must have come much later after extrusion of the main part of the basalts as all the basalts and dykes appear to be tilted to the same extent. the flexure is not a simple bending as perhaps envisaged by wager & deer (1938), as the uppermost lithosphere is too brittle; it actually consists of a multitude of normal faults, including preflexure faults (which may be ancient, reactivated faults, klausen & larsen 2002), syn-flexure faults and postflexure faults (nielsen 1975, nielsen & brooks 1981), some of which are related to other events. it is clear that the development of the flexure and intrusion of the dyke swarms took place in association with the collapse of the continental margin into the newly formed rift and were accompanied by the growth of the offshore seaward-dipping wedge of lavas (e.g. larsen, h.c. & saunders 1998). a very similar flexure is found on the west greenland margin (geoffroy et al. 1998, 1999, 2001a, b) and has been imaged on the jan mayen plateau, which is the continuation of the east greenland dyke swarm rifted off in the eocene (gudlaugsson et al. 1988). it can readily be calculated that extension along these faults accounts for at least as much dilation as that caused by the dykes (nielsen & brooks 1981). it should be noted that the work by geoffroy et al. (1998, 1999, 2001a, b) has come under scrutiny by chalmers et al. (2001) on the basis of their own studies in the area (chalmers et al. 1999a). the coastal dyke swarm has a total volume of magma that could be at least as large as that of the original overlying lavas, although it is difficult to make an accurate estimate when the dyke swarm is unknown in depth. we can gain a rough idea by calculating the integrated volume in the upper 20 km of crust, which is very conservative, but nevertheless gives volumes similar to the volume of the basalts (rucklidge et al. 1980). in addition, there is magmatic material underplated at depth, as discussed below. the dyke swarm may have formed over a significant period of time, and the impression from existing evidence is that the magmas were emplaced relatively quickly (neuhoff et al. 1997), although this cannot at present be accurately quantified. an indication of the time span involved might be obtained by a comparison of the coastal dykes cutting and cut by the skaergaard intrusion, whose age and cooling history are accurately known. superficially, the number cutting the intrusion is relatively small, but a quantitative study would be necessary as the skaergaard intrusion is placed somewhat outside the main zone of coastal dyking. early dykes, which have a composition and age indicating they were feeders to the overlying lavas, including the lower basalts (e.g. gill et al. 1988; hanghøj et al. 2003; lenoir et al. 2003; tegner et al. 2008), have been strongly tilted as shown by equivalent tilting in lavas as well as in sedimentary rocks where these occur. the early tilted dykes are cut by later swarms, which finally become vertical and have different and often more evolved compositions (e.g. hanghøj et al. 2003). it is clear that the latest dykes, intruded after formation of the coastal flexure, remain close to vertical. the late dykes constrain the length of time of flexuring to only a few million years, and lenoir et al. (2003, who also estimated a strain rate for this region) determined the duration to be c. 2.9 ma by 40ar/39ar dating of dykes at kap wandel and kap gustav holm. as previously noted inland dykes are sporadic, but, from the hinge line of the flexure, dyke abundance increases rapidly seawards and may reach >95% of the outcrop area in some areas, being effectively a sheeted dyke swarm. the coastal swarm is here similar to the sheeted dykes of ophiolites, such as the troodos massif on cyprus, and such dyke complexes are thought to make up layer 2 of the oceanic crust. just as in the sheeted dyke swarms, the east greenland dyke swarm, where it is densest, hardly has any screens of country rock. furthermore, dykes have intruded one into another leaving one-sided chilled margins, as described by kidd & cann (1974). the offshore extension of this dyke swarm is unknown although it has been discussed by larsen, h.c. (1978). segmentation of the margin a notable feature of the margin is its segmentation, with en échelon segments filled from central magma chambers (myers 1980; karson & brooks 1999; klausen & larsen 2002; callot et al. 2001, 2002). the segmentation occurs on several scales. a first-order segmentation is represented by the kangerlussuaq triple junction and the triple junction to the south of greenland where the reykjanes, mid-atlantic and labrador sea ridges meet (nielsen, t.k. et al. 2002). secondand third-order segmentation is marked by largely strike-slip faulting at a high angle to the dyke swarms, although vertical 6666 offsets are apparently important in some cases. secondorder segmentation may equate with the geophysical regions described by larsen, h.c. (1990, see fig. 19), distinguished on the grounds of morphology, differences of crustal type, subsidence history, etc. these units are spaced c. 100–120 km. the third-order accommodation zones are spaced c. 25–39 km apart and are usually marked by abrupt changes in dyke intensity, fault geometries and flexure geometries. spacing of the gabbro intrusions may be related to this segmentation, while the possibility of nephelinite-carbonatite volcanism, suggested in the previous section on nephelinites and carbonatites, may have been controlled by the secondorder accommodation zones (fig. 19). guarnieri (2011) described two periods of strike-slip faulting separated by the formation of the coastal dyke swarm in sødalen, although the significance to the overall tectonics of the margin was unclear. however, it was suggested that the third phase of faulting was connected to the regional uplift. vertical and horizontal distribution of igneous material in the crust as already noted, the igneous rocks at the surface are only a small part of the total amounts of magma involved. an important part of the magma budget is underplating of magma beneath the margin. such a process has been argued by brooks (1985a) and cox (1993) as being responsible for the high elevations of such regions, although it does not explain the late timing of the uplift (see the geomorphology section). the margin was imaged in some detail by the sigma project (see e.g. korenaga et al. 2000, 2002; holbrook et al. 2001; hopper et al. 2003), which showed thick wedges of high velocity material under the margins that was interpreted as underplated basaltic material. the wedges were found to decrease with distance to the south as did the thickness of ocean crust from c. 30 to c. 17 km (for location of profiles see fig. 1). this decrease is taken to reflect waning magma production away from the icelandic hotspot and corresponding decreasing amounts of underplated material. note, however, that white et al. (2008) claimed that the magmatic material was emplaced largely as sills rather than by underplating, and in east greenland the dyke swarms represent a volume of a similar order of magnitude as the lavas, as discussed previously. holbrook et al. (2001) further estimated that the ratio of intruded to extruded magma is 1.5:1, which broadly corresponds to this author’s estimates as it includes the underplated material and is therefore 50% higher. across the continent–ocean transition, the thickness of the crust diminishes from c. 30 to c. 8 km, which was explained e.g. by holbrook et al. (2001) as a result of vigorous active upwelling during the early stages, giving way to passive upwelling later. the transition from archaean crust to true oceanic crust takes place over c. 400 km (korenaga et al. 2000; hopper et al. 2003). summarising these results, we see that large amounts of igneous material crystallised at the base of and within the crust. also, the amount of igneous material decreases both laterally away from the presumed hotspot and temporally after initial break-up. in particular, decreasing melt volumes, both with distance from the centre and with distance from the initial break-up site, are consistent with the idea of a plume consisting of a large head and a thin tail (e.g. campbell 2007). generation of copious amounts of magma by a plume requires that the plume has excess temperature, suggested to be c. 300°c, relative to that at mid-ocean ridges, where extension as the plates move apart causes passive melting by simple decompression in the mantle (e.g. klein & langmuir 1987). note that these values are expressed as ‘potential temperature’, that is, the equivalent temperature of the magma at depth when adiabatically decompressed to surface pressure. plumes with excess temperature have thermal buoyancy and cause melting by an excess of temperature, known as active melting. whether plumes show such excess potential temperature is a long-standing controversy and surprisingly difficult to determine to everyone’s satisfaction. one of the latest discussions on this subject is that of herzberg & gazel (2009) who showed, using petrological arguments, that many large igneous provinces had potential temperatures (as defined above) of c. 300°c. according to herzberg & gazel (2009) the north atlantic province was one of the regions of high potential temperature, but in common with other provinces (e.g. the caribbean large igneous province with its present manifestation at galapagos), the temperature decreased with time to a present excess potential temperature of c. 100°c in iceland. this reinforces the long-held assertion that plume heads are the sites of most voluminous magmatism (e.g. campbell & griffiths 1990). in a modelling study of the ocean drilling program legs 152 and 163 (fig. 1), an area also covered by the sigma project (holbrook et al. 2001), it was 67 shown by armitage et al. (2008) that the basalt compositions could be reproduced by including a hot layer and with a pulse of rapid extension. no anomalous mantle was required. this scenario is nevertheless broadly consistent with existing requirements of the plume hypothesis (e.g. campbell & griffiths 1990, campbell 2007) and explains the rapid decrease in igneous production, as measured by crustal thickness, towards the ocean. however, it is also clear that a complete understanding of magma production and continental break-up has still not been acquired. plate-tectonic patterns since the formation of the first ocean floor c. 55 ma ago (magnetochron 24r), spreading appears to have been relatively simple south of present-day iceland, equivalent to the areas south of kangerlussuaq in the paleocene. in this area, the host rocks are precambrian. to the north, however, spreading was more complex. in the blosseville kyst area, larsen, l.m. & watt (1985) showed that abortive early break-up was to the west, probably along the n–s-trending zone of heavy fracturing and passing into the inner reaches of scoresby sund and perhaps farther to the caledonian main fault, as shown in figs 2 and 14a. the lower two basalt formations were, at least, partially sourced along this zone. later, the rift jumped to the east, forming the prominent bulge that is blosseville kyst (larsen, l.m. & watt 1985; larsen, l.m. et al. 1989; larsen, h.c. 1988). early spreading also took place along the ægir ridge, and extending northwards to the jan mayen fracture zone, in the area where the lithosphere had been subjected to late caledonian and mesozoic rifting events. later, from the middle eocene (chron 20) to the late oligocene at c. chron 6–7 (talwani & eldholm 1977; nunns 1983a, b; bott 1987; larsen, h.c. 1988; eldholm et al. 1990) the spreading axis gradually relocated to the west on the present kolbeinsey ridge (fig. 23a). this led to the splitting off of the jan mayen plateau, previously part of greenland (scott et al. 2005), reducing the bulge in the spreading axis, and with a consequent profound modification of the east greenland margin in these regions. the northwards continuation of the flexure and its associated dyke swarm is now to be found on the eastern side of the jan mayen plateau and was imaged by gudlaugsson et al. (1988). it may be that this event led to new magmatic episodes in east greenland (as argued by several authors, e.g. brooks et al. 2004). the balanced extinction and propagation of ridges were first described by hey et al. (1989) from the east pacific rise and applied to the reykjanes ridge by hey et al. (2010). vink (1984) showed how the greenland–iceland– faroe ridge area and the vøring plateau could be explained by a plume, moving relative to the plates, coupled with a transform fault (jan mayen) and a ridge jump (ægir to kolbeinsey). this once well-received model (fig. 23b) seems to have become forgotten by workers in the area. vertical movements and erosion another intractable problem concerns the topography of the area. this section extends the discussion of the subject introduced in the section on geomorphology, with a more regional view. east greenland constitutes a mountain range comparable in bulk to collisional mountain belts, such as the european alps, although east greenland has not experienced significant compressional tectonics since the palaeozoic, when the caledonian mountains were formed and rapidly peneplained. moreover, east greenland is not unique. the same anomalous situation is found throughout the north atlantic ocean and was called the ‘icelandic phenomenon’ by vogt (1974). thus, high margins flank the atlantic ocean in eastern canada, west greenland, norway and scotland; all areas thought to be affected by the icelandic plume. what are the timing and origin of such uplifts? topography, of course, generally reflects the thickness of the lithosphere, except where dynamic effects in the mantle cause anomalies (positive over rising mantle, e.g. plumes, and negative where mantle descends, e.g. over subduction zones). dynamic uplift has a long wavelength of hundreds to thousands of kilometres and is transient (braun 2010), so the short-wavelength and apparently relatively stable uplift seen in east greenland doubtlessly owes its origin to thickened lithosphere. this is most likely caused by gabbroic underplating or major sill emplacement at depth, as mentioned in several places here, and favoured by the results of the sigma project (holbrook et al. 2001). however, as seen already, the timing is difficult to reconcile, as was also pointed out by clift et al. (1998). 6868 k o lb ei n se y r id ge re yk ja ne s rid ge jan m ayen fracture zone greenland a b 20°w kialeeq area s c a n d in a v ia kangerlussuaq area blosseville kyst vøring plateau m o h n s r id ge 60°n jan mayen fracture zone 60°n tjörnes fracture zone jan mayen plateau 10°w 0° 10°e 20°e ægir r id ge ( ex ti nc t) 70°n mesters vig area faroe islands 40°w model observed 30°w cenozoic oceanic crust late eocene to late oligocene magmatism fracture zone ocean ridge (active or extinct) highly extended continental crust continental shelf (continental crust) rockall trough rockall-hatton bank propagator tip failing spreading axis doomed spreading axis (ægir ridge) offset zone (jan mayen plateau) propagating spreading axis (kolbeinsey ridge) failed spreading axis greenland–iceland–faroe ridge area vøring plateau continental crust continental crust hotspot track hotspot track transform fault hotspot plateau abandoned plateau channelling modified from vink (1984) 69 the complicated interplay between several factors (uplift, erosion, climate) has been succinctly covered by england & molnar (1990) and molnar & england (1990). they argue that many workers have postulated late uplift using proxies that may instead reflect global climate change. in east greenland, however, the late cenozoic uplift seems to be well established by fissiontrack studies, although its precise timing has not yet been pinpointed. as noted above, the idea that the uplift was caused by crustal underplating was attractive as it shows no signs of thermal relaxation, but this explanation is unlikely as the uplift was delayed by tens of millions of years (as noted above). medvedev et al. (2008) argued that the erosion of deep fjords and the transfer of sediments to the shelf were sufficient to account for the uplift of the high peaks, suggesting that the erosion was glacier driven. it remains to be shown that the majority of the uplift can be assigned to the pleistocene, although undoubtedly some of this uplift can be accounted for in this way, as remarked by brooks (1979 & 1985a). also, the present fjords may represent large pre-pleistocene alluvial valleys only modified to a relatively minor extent by glacial activity so that uplift had begun at an earlier stage. rather, it is tempting to view the distribution of high land round the north atlantic ocean as being somehow connected to the ‘icelandic anomaly’: the sea-floor also being unusually shallow here with iceland itself as the only place where the mid-atlantic ridge becomes subaerial. it therefore seems that the mountain uplift in east greenland (and other areas of the north atlantic region) critically depends on a consensus to explain iceland’s anomalous topography. geomorphological studies (brooks 1979) have shown that the east greenland margin was subjected to two main uplift events: the formation of a domal structure rising to c. 6 km above present sea level and centred on the kangerlussuaq area, and a regional uplift of c. 2 km giving the high east greenland coastline. these two topographic features could well have different causes. fission-track studies (gleadow & brooks 1979; hansen 1992, 1996, 2000; hansen & brooks 2002; hansen & reiners 2006) show that uplift and erosion took place c. 20 ma after the continental breakup and continued into the neogene, but precise mechanisms for these events remain obscure, the delay probably being inconsistent with the idea of underplating as suggested by brooks (1985a), cox (1993) and white et al. (2008). one of the arguments relating to uplift are the socalled v-shaped ridges on the reykjanes ridge (see hey et al. 2010 for references to the rather extensive literature). two recent papers continue the old idea that these ridges are caused by pulses of enhanced magmatism travelling along the ridge axis and causing uplift (poore et al. 2011; see also white & lovell 1997, and hartley et al. 2011 for a similar explanation for north sea events). however, hey et al. (2010) refuted the pulsing plume model and claimed that these diachronous ridges are due to ridge jumps, as was also favoured by hardarson et al. (1989). time frame for magmatism and margin deformation high-precision ages for other parts of the north atlantic province have gradually appeared (e.g. berggren et al. 1995; pearson et al. 1996; storey et al. 1998; hamilton et al. 1998; gamble et al. 1999; heister et al. 2001; chambers et al. 2005) allowing a comparison with the ages summarised here for east greenland. although space does not permit a detailed account of the present state of knowledge regarding the geochronology of east greenland and its relation to the rest of the north atlantic igneous province, fig. 24 presents a summary of the east greenland geochronological data, as discussed above. see also the compilation in the appendix. in this figure, it is important to note that some of the events are major, such as the formation of the main basalts at c. 56 ma, and other events, such as the intrusion of late alkaline dykes, are volumetrically quite minor. in general the later events were not voluminous in comparison to the main basalts. figure 24 also includes data on major tectonic adjustments in the north atlantic area facing page: fig. 23. a: plate-tectonic features of the north atlantic region. b: the model of vink (1984) for plume control of north atlantic oceanic plateaux (redrawn and simplified). model with two successive stages shown left and application of the model to the right. in its early stages the plume feeds both the vøring plateau and the faroe islands, being at equal distance (marked by arc) from the plume centre. with spreading, the vøring plateau becomes too distant due to a change in direction of plate movement, and the plume now only feeds the greenland–iceland–faroe ridge area. 7070 and estimates for magma production from storey et al. (2007b) around the break-up event. a perusal of fig. 24 does not reveal any striking patterns. the main activity was with the early plate separation in the labrador sea at 62–60 ma and the subsequent plate separation in the north-east atlantic around c24r or 57–55 ma. these two events clearly relate to the extrusion of the lower basalts and the main basalts. after this, intrusive activity continued, most notably in the kangerlussuaq area, with much later activity around kialeeq to the south (fig. 17) and mesters vig to the north (fig. 14), which corresponds to late minor activity in west greenland. as discussed previously, this may be related to rifting of the jan mayen plateau and subsequent cessation of spreading on the ægir ridge, although clear synchronicity is not obvious. the cause of the latest miocene activity at vindtoppen is also obscure, although storey et al. (2004) concluded that this was due to partial melting of metasomatised mantle, perhaps triggered by an uplift event that seems to have occurred at this time (i.e. 13–14 ma). the intrusions at kialeeq seem to correlate with major plate-tectonic adjustments as suggested in fig. 23. of particular interest here, is the sanidine-bearing tuff whose temporal position is shown in fig. 3. as previously mentioned, this tuff is thought to be identifiable over a wide north atlantic area, both in the ash layers in northern denmark, in those of the north sea, in eastern england and in dsdp hole 550, 14°14´w, 48°25́ n, south-west of ireland, although heister et al. (2001) and larsen, l.m. et al. (2003) do not concur on its source being necessarily the gardiner complex. related ash layers, perhaps some of the largest on record, even extend to austria. such layers are therefore important time markers, especially as in hole 550 it was deposited shortly after the paleocene–eocene thermal maximum, a subject that has been under intense debate in recent years and the cause of which is still not well understood. one previously mentioned hypothesis is that of svensen et al. (2004, 2010), who argued that the required amounts of methane were generated by sill intrusion into carbonaceous sedimentary rocks in the north atlantic igneous province. this is appealing as the timing is correct. the age of the paleocene– eocene thermal maximum has recently been greatly improved, first by the work of westerhold et al. (2007, 2009). 40ar/39ar dates on the sanidine are many, but calibration uncertainties in the 40ar/39ar method mean there is a relatively large uncertainty in the results. probably the best estimate to date is a u/pb age on zircon from a section on spitsbergen (svalbard archipelago, fig. 1) giving 55.785 ± 0.034 ma (charles et al. 2011). an attempt should be made to date the perovskites that occur in the ashes of the fur formation of denmark. this, combined with astronomical tuning of the sedimentary rocks, is of great importance, not only for correlations throughout the north atlantic region, but also for calibration of the geological time scale and our understanding of important environmental events such as the paleocene–eocene thermal maximum. mechanisms of north atlantic rifting another important question, touched on previously, is the apparent asymmetry of the north atlantic and norwegian–greenland seas. asymmetric features include: generally higher topography on the greenland side (indeed the rockall-hatton bank and vøring plateau are submerged), larger volumes of igneous rocks (according to hopper et al. 2003, nearly twice the amount on the greenland side), different magmatic compositions (syenites and nephelinites are rare on the european side and carbonatites have not been reportfacing page: fig. 24. summary of the geochronology of the north atlantic igneous province. timings shown are based on a compilation of ages deemed to be the most reliable: 40ar/39ar ages that fulfil accepted criteria, u-pb ages and occasional re-os ages (see also appendix). plate-tectonic events shown on the left are based on references in the main text. magma productivity around the time of break-up from storey et al. (2007b). time scale from gradstein et al. (2005). data sources: british hebridean province: archer et al. (2005), chambers & fitton (2000), chambers & pringle (1999, 2001), chambers et al. (2005), gamble et al. (1999), gibson et al. (1995), hamilton et al. (1998, 1999), hansen, d.m. (2006), pearson et al. (1996), troll et al. (2008). faroe islands: storey et al. (2007b). east greenland south of scoresby sund and offshore: hamilton & brooks (2004), hansen et al. (2002), gleadow & brooks (1979), heister et al. (2001), peate, d.w. et al. (2003), sinton & duncan (1998), storey et al. (2004, 2007b), tegner & duncan (1999), tegner et al. (1998b, 2008), wotzlau et al. (2011). east greenland north of scoresby sund: brooks et al. (1979, 2004), price et al. (1992), rex et al. (1979), upton et al. (1995) and unpublished data by c. tegner and c.k. brooks from the mesters vig area. west greenland: storey et al. (1998). 71 cessation of spreading on ægir ridge cessation of spreading in labrador sea. spreading change in lena trough first rifting of jan mayen plateau break-up between greenland and europe break-up in labrador sea a p p ro x im at e ti m in g o f m ai n p la te -t ec to n ic ev en ts b ri ti sh h eb ri d ea n p ro vi n ce si ll in tr u si o n i n s ed im en ta ry b as in s (a p p ro x im at e) fa ro e la va s im ili k ga b b ro ( ap p ro x .) o ff sh o re l av as ( o d p ) e as t g re en la n d l av as w es t g re en la n d l av as k an ge rl u ss u aq i n tr u si o n s w es t g re en la n d la m p ro p hy re s et c. n o rt h er n e as t g re en la n d in tr u si o n s an d l av as k ia lin eq f el si c in tr u si o n s ? malmbjerg (porphyry molybdenum) ? small isles youngest mourne mt. granites red hills granites skye ? 65 60 55 50 45 40 35 30 25 ma o lig oc en e eo ce ne pa le oc en e la te ea rl y la te m id dl e ea rl y la te m id dl e ea rl y biotite granite kærven, skaergaard intrusions prinsen af wales lavas miocene basalts 13 ma 63–64 ma (archer et al. 2005) c ha tt ia n r up el ia n pr ia bo ni an ba rt on ia n lu te tia n y pr es ia n t ha ne tia n se la nd ia n d an ia n cretaceous neogene epoch age chron polarity c6b c6c c7 c7a c8 c9 c10 c11 c12 c15 c16 c17 c18 c19 c20 c21 c22 c23 c24 c25 c26 c27 c28 c29 c30 flammefjeld (porphyry molybdenum) youngest dykes main basalts lower basalts sp re ad in g on ly a lo ng k ol be in se y r id ge sp re ad in g al on g k ol be in se y an d æ gi r r id ge s sp re ad in g al on g æ gi r r id ge ex te ns io n in n or th a tla nt ic ri ft s ys te m post-break-up break-up-related hiatus pre-break-up 1 0 0 0 2 0 0 0 3 0 0 0 km3/km/ma magma productivity blosseville kyst – faroe islands sector (storey et al. 2007b) 7272 ed) and a much longer period of igneous activity in east greenland. the asymmetry has important bearings on the mechanism of break-up, the loci of magmatic activity and the shapes of sedimentary basins. it is also a deeply controversial subject. continental break-up is usually depicted as simple ductile necking (pure shear) of the lithosphere. however, rosendahl (1987) showed the asymmetry of the african rift basins, typically half-grabens, suggesting that rifting takes place by detachment (simple shear), an idea applied to many rifts worldwide by a range of authors. for example, wernicke & tilke (1989) used this model for the central atlantic margin of the united states and pointed out that the woodlark basin rift of papua new guinea advances by a series of rift basins forming along low angle detachment faults (speckbacher et al. 2011 and references therein). the simple shear model has been particularly popular with norwegian geologists, explaining their own margin, where, it should be remembered, large amounts of observational material have been gathered in the search for hydrocarbons. torske & presvik (1991) attributed the nephelinites of the vestbronna formation on the møre platform (fig. 1) to simple shear, while mosar et al. (2002a) regarded the norwegian margin as defined by low-angle detachment faults, which change their polarity across the jan mayen fracture zone. the møre margin is thus a lowerplate margin, and the vøring–trøndelag margin an upper-plate configuration. consequently, the central fjord district of east greenland including the basalt plateaux would be lower-plate, while the liverpool land margin would be upper-plate. to extend this southwards, the margin from scoresby sund to kangerlussuaq would again be lower-plate and south of this upper-plate. could this explain some of the geological observations discussed in this paper; for example the distribution and preservation of the basalt plateaux? certainly, the postulated changing polarities appear to take place along some of the accommodation zones described by karson & brooks (1999) or along changes of crustal character found by larsen, h.c. (1990), as shown in fig. 19. on the other hand, it is worthwhile noting that studies of the east pacific rise have shown that it too is asymmetric (melt seismic team 1998; levi 1998), although there is no suggestion that simple shear is the controlling mechanism here. rather it is caused by westward drift of the underlying upwelling mantle. nevertheless, hinz et al. (1987) stressed the symmetry of the norwegian and east greenland margins. however, the object of their studies was primarily the seaward-dipping reflectors, which may disguise an earlier asymmetry, and more recent work favours an asymmetric structure (hopper et al. 2003). summary and perspectives some of the most important conclusions regarding the evolution of the east greenland rifted volcanic margin and the adjacent north atlantic region can be summarised as follows: (1) igneous activity within the north atlantic province was widespread from its earliest inception (as noted by several authors, e.g. chalmers et al. 1995) covering an area c. 2000 km in diameter from baffin island and west greenland to the british isles. the igneous activity commenced c. 64 ma ago, judging by the age of sills in the rockall trough (archer et al. 2005), and seems to have become rapidly contemporaneous over this entire area. in east greenland, this phase is represented by the lower basalts at kangerlussuaq, which are more picritic, more contaminated and more explosive than the succeeding main basalts. the most voluminous burst of volcanic activity was at c. 56–55 ma and occurred in east greenland, represented by the main basalts which cover c. 65 000 km2 and have a thickness >7 km. (2) after c. 55 ma, major extrusive activity in east greenland had ceased. however, activity continued in the mesters vig area until c. 25 ma, and extensive felsic activity occurred in the kialeeq area at c. 35 ma (fig. 16). similarly, minor activity 73 occurred sporadically in west greenland until at c. 30 ma (storey et al. 1998). further, renewed basaltic volcanism is recorded in the miocene vindtoppen formation (c. 14 ma, storey et al. 2004), and up to the present time at jan mayen (trønnes et al. 1999), the vesteris bank (haase & devey 1994) and the icelandic flank volcanoes such as snæfellsjökull. (3) volcanism continues today in iceland and can be regarded as a direct extension of the events occurring in east greenland in the palaeogene. the postulated mantle plume is now supposed to be situated under vatnajökull (wolfe et al. 1997). the main difference from east greenland is that continental crust is currently not thought to be present in iceland. this probably accounts for much of the contrast between palaeogene greenlandic volcanism and recent icelandic volcanism such as the abundance of syenites in east greenland and the lack of strongly undersaturated rocks of the nephelinite-melilite clan in iceland, which may be the result of a thicker lithospheric lid in east greenland and be caused by the presence of continental crust. contamination of icelandic rocks with amphiboliteand granulite-facies gneisses, which is observed in east greenland, would not be expected and has not been observed in iceland, although felsic magmas are relatively abundant. it is worth noting that the earliest rifting event discernable in iceland took place from 25 to 15 ma ago. volcanic cycles in iceland are connected to ridge jumping, which is thought to occur in order to bring the spreading centre over the plume head. the rifting events in iceland had an approximate life span of 8–12 ma (exact values are difficult to decipher as there is an overlapping period of propagation and dying: hardarson et al. 1997; pringle & hardarson 2000). this is similar to the 5–10 ma clustering of activity in the different areas of east greenland (fig. 24). doubtlessly, these events are closely connected with the plate-tectonic development of the northernmost atlantic region, although uncertainties in the timing of plate-tectonic events do not at present allow a precise correlation. in iceland, the clustering seems related to the movement of the plate boundary relative to the plume, with consequent adjustments of the former. (4) numerous attempts have been made to relate magmatic activity to the passage of the supposed icelandic plume. the most widely accepted scenario is as follows. the earliest activity beginning c. 62 ma ago is thought to reflect the arrival of the plume in the area, although the passage of the plume depends on whether one relies on surface manifestations (brooks 1973; forsyth et al. 1986; white & mckenzie 1989) or models based on palaeomagnetism (lawver & müller 1994), as previously discussed in the introduction to the section on the main basalts. later activity at 55 ma reflects the first successful break-up between the faroe islands and greenland, while activity in the range 52–47 ma is thought to mark the passage of the plume stem under east greenland (tegner et al. 1998b). the cause of later igneous activity, largely at 35 and 25 ma in the kialeeq and mesters vig areas, remains obscure, but may be related in part to plate reorganisation in the north atlantic ocean, such as the demise of spreading in the labrador sea and of the ægir ridge, the splitting off of the jan mayen plateau and a change in spreading along the lena trough, all taking place around the time of magnetochron c13, dated at c. 33 ma (fig. 24). the ultimate cause of the formation of the north atlantic igneous province, i.e. plume or non-plume, has been discussed by numerous other authors (see e.g. depaolo & manga 2003; foulger et al. 2005; campbell 2007; foulger 2010) and lengthy further discussion is out of place here. nevertheless, our understanding of this major geological event, so beautifully displayed in east greenland, seems to be far from reaching a consensus, either on this subject or others such as the cause and timing of uplift, and brief comments are warranted. the present discussion is confined to a summary of the main features that are believed to be strong evidence for a plume origin as originally suggested (brooks 1973). these include the rapid eruption of high-temperature magmas over a wide area. although controversial, most petrologists regard picrites as evidence of high-temperature source regions. regional uplift of the lithosphere is evidence for the arrival of the temperature anomaly under the area. extraordinarily large volumes of magma, far in excess of that which can be produced by simple extension, were produced at the height of the break-up. the subsequent decrease in magma volumes was pre7474 dicted in the original exposition of impacting mantle plumes by campbell & griffiths (1990) and recently modelled by armitage et al. (2008) and herzberg & gazel (2009). the trace of the supposed plume trail in the aseismic greenland–iceland–faroe ridge area (e.g. wilson 1963; brooks 1973; vink 1984; richards et al. 1989) was noted many years ago. in addition, a decrease in productivity away from the supposed plume centre and orthogonal to the plume trace, as shown by the sigma transects (fig. 1; holbrook et al. 2001; hopper et al. 2003). even the double pulse represented by the lower basalts and the main basalts may be modelled (shu-chuan lin & van keken 2005). finally, seismic imaging of the hot plume has been claimed for both iceland and hawaii (wolfe et al. 1997, 2009). additional evidence for this is provided by the discovery that the west greenland and baffin island lavas (and by extension those of the rest of the province, as well as other likely plume sites worldwide) are derived from regions of primitive mantle. they seem to coincide with regions of anomalous seismic behaviour, known as the large-low-shear-wave-velocity provinces (torsvik et al. 2010; jackson & carlson 2011). concluding remarks the east greenland margin, with its well-exposed and well-documented features, is an incomparable laboratory for the study of the interplay between mantle flow, lithospheric tectonics and the formation of a rich diversity of magmas. it is thus an outstanding test area for the existence of a mantle plume. in the above, the extreme complexity of the east greenland margin has been described. it is both typical and better exposed than most comparable places such as brazil, south africa or india, not to mention submerged large igneous provinces such as ontong java. these complexities have to be taken into account in any comprehensive explanation of continental rifting leading to ocean formation, making a detailed documentation of the magmatic, tectonic and chronological evolution of the margin of paramount importance. however, it must be concluded that not only the fundamental cause of large igneous provinces is still in debate but lesser questions are far from being understood. many basic processes remain enigmatic – as demonstrated by the skaergaard intrusion, one of the best studied intrusions in the world, but nevertheless still highly controversial. during the time this author has worked in east greenland, a wealth of new observations have been made: new intrusions discovered, tectonic events revealed, refinement of the geochronology, knowledge of the offshore extensions gained and large systematic collections of both samples and data assembled. this makes it clear that observation plays a crucial role and theory must come in second place. many earlier models have been shown to be inappropriate when new discoveries have been made in the field or new collections have been studied. additionally many elegant models have been proposed on the flimsiest of evidence. it is obvious that earlier ideas were often based on too little observation and too few samples. this is not necessarily a criticism: it was often not possible to do better, either because of the difficulties of field work or by the limitations imposed by laboratory procedures. today, with the common availability of stol aircraft and helicopters, the only barrier to transportation and hence observation is the cost. furthermore, the great improvements that have come about in analytical and geochronological methodology; in accuracy, precision, speed and scope have already shown results. of all improvements, micro-beam methods may be the most significant as they allow single-gain dating of minerals with high precision as well as detailed study of minerals with a view to tracking the history of magmatic development. similar advances in seismic methods and in the interpretation of fission-track data have also been made. although the north atlantic igneous province, especially the hebridean part, is probably the most thoroughly researched province of its type, many important details still continue to emerge. not least, the detailed geology of the offshore areas is still only known in outline, although much information is emerging from the oil and gas industry west of the shetland islands. until we come close to a precise knowledge of what all parts of the province contain and how the province developed, sophisticated models relating the magmatism to plate-tectonic events and plume properties will remain ill-founded. 75 acknowledgements professors l.r. wager and w.a. deer introduced me to east greenland in my student days and inspired me with the mysteries of igneous rocks and their minerals. i would like to thank my students and colleagues who contributed to the kangerlussuaq project, in particular: jens-christian andersen, stefan bernstein, karen hanghøj, rune larsen, peter momme, troels f.d. nielsen, morten riishuus and christian tegner. dennis k. bird has repeatedly provided invaluable scientific input. our research was financed by statens naturvidenskabelige forskningsråd (danish natural science research council) over many years and supported in the later years by robert gannicott, platinova resources limited / harry winston diamond corporation. work continued in the 1990s under the auspices of the now defunct danish lithosphere centre, financed by the danish national research foundation. the considerable contribution by the centre has been made plain in this publication. reviews by dennis k. bird, stanford university, and christian tegner, aarhus university, have resulted in numerous detailed and valuable amendments to the manuscript. i am also very grateful to my inuit friends, particularly the late ulrik pîvât and his son abia and the uitsatikitseq family, all of kuummiut, for hospitality and transport over the years. maps shown in figs 9 and 10 are compiled from unpublished maps by dennis k. bird, stanford university, christian tegner, aarhus university, and stefan bernstein and nicholas rose (now avannaa resources ltd, copenhagen) made in 1987, 1988, 1989 and 1990 and funded by the danish lithosphere centre and nsf-ear grants no. 84-18129, 86-06256, 88-03754 and 90-06962. 7676 references geochronology. chemical geology 200, 241–253, http://dx.doi. org/10.1016/s0009-2541(03)00202-x barker, a.k., baker, j.a. & peate, d.w. 2006: interaction of the rifting east greenland margin with a zoned ancestral mantle iceland plume. geology 34, 481–484, http://dx.doi.org/10.1130/ g22366.1 bearth, p. 1959: on the alkali massif of the werner bjerge in east greenland. meddelelser om grønland 153, 63 pp. beckinsale, r.d., brooks, c.k. & rex, d.c. 1970: k-ar ages for the tertiary of east greenland. bulletin of the geological society of denmark 20, 27–37. berggren, w.a., kent, d.v., swisher iii, c.c. & aubrey, m.-p. 1995: a revised cenozoic geochronology and chronostratigraphy. in: berggren, w.a., kent, d.v. aubrey, m.-p. & hardenbol, j. 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(1998b) 40ar/39ar 56.2 0.6 1σ * plateau imilik gabbro iii, wehrlite sill 40ar/39ar 416822 mica 49.3 0.6 1σ isochron 40ar/39ar 49.2 0.2 1σ * plateau imilik gabbro iii, wehrlite sill 40ar/39ar 416804 mica 49.8 1.2 1σ * isochron 40ar/39ar 52.1 0.6 1σ plateau kruuse fjord, gabbro pegmatite 40ar/39ar 428702 mica 48.0 1.2 1σ * isochron 40ar/39ar 51.5 0.3 1σ plateau igtutarajik, gabbro pegmatite 40ar/39ar i-7 mica 47.0 0.3 1σ isochron 40ar/39ar 47.0 0.2 1σ * plateau kap edvard holm, mls wehrlite sill 40ar/39ar pct-75 mica 48.0 0.8 1σ isochron 40ar/39ar 47.3 0.3 1σ * plateau kap edvard holm, lls gabbro pegmatite 40ar/39ar keh302 amphibole 48.5 0.4 1σ isochron same sample as nevle et al. 1994) 40ar/39ar 48.8 0.2 1σ * plateau lilloise intrusion, syenite vein 40ar/39ar ckb71-54c amphibole 50.2 1.0 1σ isochron 40ar/39ar 50.0 0.4 1σ * plateau sorgenfri sills, i.c. jacobsen fjord dolerite 40ar/39ar 40006 whole-rock 54.7 0.9 1σ isochron 40ar/39ar 56.3 0.9 1σ * plateau sorgenfri sills, fairytale valley dolerite 40ar/39ar 413907 plagioclase 56.2 1.1 1σ isochron 40ar/39ar 56.0 0.4 1σ * plateau nuuk diorite 40ar/39ar 312001 biotite 36.2 0.6 2σ * isochron tegner et al. (2008) ikâsangmît syenite intrusion 40ar/39ar 429360 amphibole 37.2 2.9 2σ * isochron patulajivit gabbro complex 40ar/39ar p-175 biotite 47.2 0.8 2σ total fusion søndre aputitêq, gabbro-granite 40ar/39ar sa-1 amphibole 49.1 1.1 2σ * plateau gardiner complex, ijolite dyke 40ar/39ar 29904 biotite 56.1 0.1 2σ * isochron gardiner complex, melilitolite ring dyke 40ar/39ar 303825 biotite 54.3 1.2 2σ * isochron from heister et al. (2001) kangerlussuaq, biotite granite 40ar/39ar 333132 biotite 46.3 1.0 2σ * isochron from heister et al. (2001) kraemer island, syenite 40ar/39ar nm27025 biotite 50.1 1.1 2σ * isochron kangerdlugssuaq intrusion, syenite 40ar/39ar eg4583 biotite 50.8 1.1 2σ * isochron snout series syenite 40ar/39ar 312139 biotite 45.3 0.3 2σ * isochron snout series syenite 40ar/39ar 81–48 biotite 46.7 0.5 2σ * isochron kærven complex, gabbro 40ar/39ar 85659 biotite 54.8 1.4 2σ * isochron kærven complex, alkali granite 40ar/39ar 40160 amphibole 52.8 1.3 2σ * isochron wiedemann fjord, lamprophyre dyke 40ar/39ar 429285 amphibole 41.5 0.9 2σ * plateau ejnar mikkelsen fjeld, intrusion gabbro 40ar/39ar 457127 biotite 51.8 1.9 2σ * isochron borgtinderne intrusion, syenite 40ar/39ar nm5133 amphibole 46.7 1.0 2σ * plateau borgtinderne, basanite dyke 40ar/39ar 457163 biotite 36.6 0.5 2σ * isochron kronborg gletscher, nephelinite diatreme 40ar/39ar 436094 biotite 50.2 1.2 2σ * isochron location and rock type method sample¢ material age (ma) ± (ma) comment reference appendix. compilation of selected ages from east greenland† † this list attempts to include only the more reliable ages. k-ar and fission track age data omitted, and rb-sr data largely omitted due to large uncertainties. the reader is referred to the original sources for details of standards used etc. ¢ where no other letters are shown, the 6-digit sample number is prefixed by ggu (geological survey of greenland). * these ages preferred by the authors (choice between plateau and isochron ages). 9494 kangerdlugssuaq intrusion rb-sr whole-rock 50.0 1.9 2σ isochron pankhurst et al. (1976) kangerdlugssuaq intrusion pulaskite rb-sr mineral 49.9 1.0 2σ isochron jameson land, basaltic dyke 40ar/39ar 407206 plagioclase 52.6 1.4 1σ * isochron hald & tegner (2000) jameson land, basaltic sill 40ar/39ar 403031 plagioclase 52.0 1.2 1σ * isochron malmbjerg re-os malm-1 molybdenite 25.8 0.1 2σ brooks et al. (2004) flammefjeld re-os 8215128 molybdenite 39.7 0.2 2σ malmbjerg 40ar/39ar 467020 biotite 25.7 0.1 1σ * plateau skaergaard intrusion 40ar/39ar sg61 biotite 55.4 0.0 2σ plateau hirschman et al. (1997) 40ar/39ar 55.4 0.7 2σ isochron skaergaard intrusion 40ar/39ar sg61 hornblende 55.4 0.2 2σ plateau 40ar/39ar 55.5 0.8 2σ isochron skaergaard intrusion u-pb zircon 55.59 0.1 2σ hamilton & brooks (2004) nansen fjord fm, lava 40ar/39ar 426296 plagioclase 57.7 0.5 2σ plateau storey et al. (2007b) 40ar/39ar 57.5 0.6 2σ isochron nansen fjord fm, lava 40ar/39ar 194021 plagioclase 59.2 1.4 2σ plateau 40ar/39ar 56.8 2.3 2σ isochron skjoldungen, dyke (fig. 1) 40ar/39ar 940211 plagioclase 61.4 1.0 2σ plateau 40ar/39ar 61.1 1.6 2σ isochron tugtilik, carbonatite dyke 40ar/39ar 417150 biotite 58.3 0.9 2σ plateau 40ar/39ar 58.1 1.0 2σ isochron tugtilik, carbonatite dyke 40ar/39ar 417151 biotite 58.3 0.9 2σ plateau 40ar/39ar 57.5 1.2 2σ isochron skrænterne fm (top), lava 40ar/39ar 412251 whole-rock 55.1 0.5 2σ plateau 40ar/39ar 55.1 0.8 2σ isochron skrænterne fm, lava 40ar/39ar 421564 sanidine 55.0 0.4 2σ plateau 40ar/39ar 53.3 0.9 2σ isochron skrænterne fm, lava 40ar/39ar 436194 plagioclase 54.9 0.9 2σ plateau 40ar/39ar 56.1 4.5 2σ isochron rømer fjord fm, lava 40ar/39ar 436056 plagioclase 55.1 0.5 2σ plateau 40ar/39ar 55.1 1.6 2σ isochron rømer fjord fm, lava 40ar/39ar 421522 plagioclase 55.4 0.5 2σ plateau 40ar/39ar 55.5 1.6 2σ isochron milne land fm, lava 40ar/39ar 404107 plagioclase 56.1 0.5 2σ plateau 40ar/39ar 56.1 0.7 2σ isochron urbjerget fm, lava (fig. 2) 40ar/39ar 436409 plagioclase 61.59 1.3 1σ plateau hansen (2002) 40ar/39ar 61.19 1.5 1σ isochron urbjerget fm, lava 40ar/39ar 436411 whole-rock 60.97 1.1 1σ plateau 40ar/39ar 59.58 1.7 1σ isochron milne land fm, lava 40ar/39ar 436412 whole-rock 56.72 1.0 1σ plateau milne land fm, lava 40ar/39ar 436430 whole-rock 56.74 1.1 1σ plateau 40ar/39ar 56.93 2.4 1σ isochron milne land fm, lava 40ar/39ar 436449 plagioclase 56.70 1.7 1σ plateau 40ar/39ar 56.36 1.8 1σ isochron milne land fm, lava 40ar/39ar 436458 plagioclase 54.93 1.6 1σ plateau 40ar/39ar 54.79 2.7 1σ isochron prinsen af wales fm, pegmatite 40ar/39ar 436231 plagioclase 54.88 1.0 1σ plateau peate et al. (2003) 40ar/39ar 55.07 1.1 1σ isochron lindsay nunatak, lava (fig. 2) 40ar/39ar 436223 whole-rock 53.08 1.2 1σ plateau 40ar/39ar 52.52 1.1 1σ isochron location and rock type method sample¢ material age (ma) ± (ma) comment reference appendix. compilation of selected ages from east greenland (continued) ¢ where no other letters are shown, the 6-digit sample number is prefixed by ggu (geological survey of greenland). * these ages preferred by the authors (choice between plateau and isochron ages). 95 kap brewster, lava (all from same flow) 40ar/39ar eg7147 whole-rock 54.1 4.2 2σ plateau hansen et al. (1995) 40ar/39ar 55.0 16.0 2σ isochron 40ar/39ar eg7148 whole-rock 57.8 4.5 2σ plateau 40ar/39ar 51.0 9.0 2σ isochron 40ar/39ar eg7149 whole-rock 56.9 5.2 2σ plateau 40ar/39ar 56.0 10.0 2σ isochron 40ar/39ar eg7150 whole-rock 57.7 3.4 2σ plateau 40ar/39ar 54.0 6.0 2σ isochron 40ar/39ar eg7151 whole-rock 54.5 4.1 2σ plateau 40ar/39ar 55.0 8.0 2σ isochron kap brewster lava: preferred age 40ar/39ar 56.7 4.3 skrænterne fm, sanidine-bearing tuff 40ar/39ar 421564b sanidine 55.12 0.1 2σ mean (n=15) storey et al. (2007a) skrænterne fm, sanidine-bearing tuff 40ar/39ar 421564b sanidine 55.80 0.5 preferred heister et al. (2001) traill ø, alkaline basaltic dyke 40ar/39ar 4067 whole-rock 36.14 0.1 1σ plateau price et al. (1997) 40ar/39ar 36.87 0.2 1σ isochron 40ar/39ar 37.14 0.2 1σ plateau traill ø, alkaline basaltic dyke 40ar/39ar 4077 whole-rock 36.68 0.3 1σ plateau 40ar/39ar 36.23 0.3 1σ isochron 40ar/39ar 36.17 0.3 1σ plateau traill ø, tholeiitic basaltic dyke 40ar/39ar 4058 whole-rock 52.00 2.0 1σ isochron 40ar/39ar 56.00 5.0 1σ isochron traill ø, tholeiitic basaltic dyke 40ar/39ar 4021 whole-rock 54.45 0.4 1σ plateau 40ar/39ar 52.75 1.1 1σ plateau traill ø, tholeiitic basaltic dyke 40ar/39ar 4105 whole-rock 53.50 1.5 1σ plateau 40ar/39ar 55.80 0.9 1σ plateau 40ar/39ar 53.20 1.2 1σ isochron myggbukta complex, picrite sheet 40ar/39ar 228139 whole-rock 32.7 1.2 1σ plateau upton et al. (1995)¶ 40ar/39ar 32.0 5.0 1σ isochron hold with hope, upper series dyke 40ar/39ar 228045 whole-rock 56.6 1.9 1σ plateau 40ar/39ar 57.0 4.0 1σ isochron hold with hope, upper series dyke 40ar/39ar 228058 whole-rock 56.7 0.7 1σ plateau 40ar/39ar 51.0 6.0 1σ isochron hold with hope, upper series anka 40ar/39ar 227959 whole-rock 53.4 3.9 1σ plateau ramitic lava 40ar/39ar 53.0 6.0 1σ isochron hold with hope, lower series nephe40ar/39ar 227933 whole-rock 58.7 1.4 1σ plateau linitic lava 40ar/39ar 60.0 12.0 1σ isochron odp leg 152-917a-23r-1a (basalt) 40ar/39ar 31–36 m whole-rock 60.1 0.8 1σ plateau sinton & duncan (1998) 40ar/39ar 59.4 2.4 1σ isochron odp leg 152-917a-23r-1a (basalt) 40ar/39ar 31–36 m whole-rock 60.8 1.0 1σ plateau 40ar/39ar 60.6 1.8 1σ isochron odp leg 152-917a-24r-1 40ar/39ar 8–10 m feldspar 62.2 0.4 1σ plateau 40ar/39ar 62.3 1.4 1σ isochron odp leg 152-917a-40r-4 (basalt) 40ar/39ar 15–20 m whole-rock 60.7 1.2 1σ plateau 40ar/39ar 60.3 4.2 1σ isochron odp leg 152-917a-52r-1 (dacite) 40ar/39ar 45–49 m whole-rock 61.4 0.6 1σ plateau 40ar/39ar 61.0 1.2 1σ isochron odp leg 152-917a-69r-2 40ar/39ar 43–53 m plagioclase 60.4 0.7 1σ plateau 40ar/39ar 60.3 1.6 1σ isochron odp leg 152-917a-83r-2 (basalt) 40ar/39ar 113–118 m whole-rock 68.2 1.6 1σ plateau 40ar/39ar 63.3 1.2 1σ isochron odp leg 152-917a-83r-3 (basalt) 40ar/39ar 46–50 m whole-rock 64.1 2.2 1σ plateau 40ar/39ar 63.3 1.0 1σ isochron odp leg 152-918d-94r-2 (basalt) 40ar/39ar 42–47 cm whole-rock 51.9 0.8 1σ plateau 40ar/39ar 51.9 1.2 1σ isochron location and rock type method sample¢ material age (ma) ± (ma) comment reference appendix. compilation of selected ages from east greenland (continued) ¢ where no other letters are shown, the 6-digit sample number is prefixed by ggu (geological survey of greenland). * these ages preferred by the authors (choice between plateau and isochron ages). ¶ information very sparse and with discrepancies between text and table, e.g. sample 228058 has ± 7.0 ma in table and ± 0.7 ma in text. 9696 odp leg 163-988a-3r-2 40ar/39ar 70–76 cm plagioclase 49.4 0.3 1σ plateau tegner & duncan (1999) 40ar/39ar 47.5 1.8 1σ isochron odp leg 163-988a-3r-3 (basalt) 40ar/39ar 17–22 cm whole-rock 49.4 0.5 1σ plateau 40ar/39ar 49.8 1.2 1σ isochron odp leg 163-988a-3r-3 (basalt) 40ar/39ar 17–22 cm whole-rock 51.1 1.0 1σ plateau 40ar/39ar 50.0 2.2 1σ isochron odp leg 163-988a-3r-3 (basalt) 40ar/39ar 17–22 cm whole-rock 49.7 0.3 1σ plateau 40ar/39ar 47.3 1.6 1σ isochron odp leg 163-989b-4r-2 (basalt) 40ar/39ar 102–110 cm whole-rock 57.1 1.4 1σ plateau 40ar/39ar 54.1 1.2 1σ isochron odp leg 163-989b-5r-2 (basalt) 40ar/39ar 28–33 cm whole-rock 57.1 1.4 1σ plateau 40ar/39ar 57.9 3.8 1σ isochron odp leg 163-990a-3r-1 (basalt) 40ar/39ar 56–60 cm whole-rock 56.2 1.3 1σ plateau 40ar/39ar 54.8 2.2 1σ isochron odp leg 163-990a-7r-2 (basalt) 40ar/39ar 50–56 cm whole-rock 55.4 0.7 1σ plateau 40ar/39ar 55.1 0.7 1σ isochron odp leg 163-990a-16r-4 (basalt) 40ar/39ar 7–11 cm whole-rock 56.0 0.9 1σ plateau 40ar/39ar 55.0 1.0 1σ isochron odp leg 163-990a-17r-3 (basalt) 40ar/39ar 104–109 cm whole-rock 55.5 2.2 1σ plateau 40ar/39ar 54.8 2.3 1σ isochron odp leg 163-990a-19r-5 (basalt) 40ar/39ar 21–26 cm whole-rock 55.2 1.4 1σ plateau 40ar/39ar 57.1 1.7 1σ isochron odp leg 152-917a-24r-1 (tuff) 40ar/39ar 1–3 cm feldspar 66.9 0.1 2σ isochron werner et al. (1998) odp leg 152-917a-23r-3 (lapillistone) 40ar/39ar 14–17 cm feldspar 60.6 0.2 2σ isochron odp leg 152-917a-53r-1 40ar/39ar 0–115 cm feldspar 60.6 0.8 2σ isochron (dacitic pyroclastic flow) odp leg 152-917a-23r-2 (tuff) 40ar/39ar 127–131 cm feldspar 60.6 0.2 2σ isochron kap edvard holm (keh) (peridotite) 40ar/39ar keh406b hornblende 49.4 0.3 1σ isochron nevle et al. (1994) keh (pegmatite) 40ar/39ar keh402a hornblende 49.7 0.3 1σ isochron keh (pegmatite pod in gabbro xenolith) 40ar/39ar keh274f hornblende 49.7 0.4 1σ isochron keh (pegmatite pod in gabbro xenolith) 40ar/39ar keh302 hornblende 49.4 0.3 1σ isochron keh (pegmatite pod in gabbro xenolith) 40ar/39ar keh302 phlogopite 48.6 0.2 1σ isochron keh (pegmatite pod in gabbro xenolith) 40ar/39ar keh306a phlogopite 48.6 0.2 1σ isochron keh (pegmatite pod in gabbro xenolith) 40ar/39ar k3e hornblende 42.1 0.5 1σ isochron keh (syenite) 40ar/39ar keh200 edenite 44.3 0.9 1σ isochron nordre aputitêq, miarolitic cavity in gabbro 40ar/39ar a14a phlogopite 48.1 0.2 1σ isochron location and rock type method sample material age (ma) ± (ma) comment reference appendix. compilation of selected ages from east greenland (continued) ineson et al. 2024: geus bulletin 57. 8375. https://doi.org/10.34194/geusb.v57.8375 1 of 20 research article lithostratigraphy of the portfjeld group (ediacaran – lowermost cambrian) of north greenland jon r. ineson*1 , john s. peel2 , sebastian willman2 , elias j. rugen3 , martin sønderholm4 , peter frykman5 1department for geophysics and sedimentary basins, geological survey of denmark and greenland (geus), copenhagen, denmark; 2department of earth sciences (palaeobiology), uppsala university, uppsala, sweden; 3department of earth sciences, university college london, london, uk; 4geological data centre, geological survey of denmark and greenland (geus), copenhagen, denmark; 5department for geo-energy and storage, geological survey of denmark and greenland (geus), copenhagen, denmark abstract the portfjeld formation, as originally defined, is the lowermost lithostratigraphic unit of the mainly lower palaeozoic franklinian basin in southern peary land, central north greenland. the unit crops out semi-continuously from valdemar glückstadt land to nordenskiöld fjord but is also recognised locally in northern parts of peary land and wulff land (western north greenland). regionally, it provides a key record of the early, pre-break-up history of the basin. the type succession in southern peary land has thus been the focus of recent biostratigraphical, sedimentological and geochemical study. this has demonstrated the composite nature of the unit, a lower interval (c. 190 m thick) of carbonate ramp deposits of neoproterozoic (late ediacaran) age being overlain at a karstic unconformity by a shallow marine, mixed carbonate–siliciclastic interval (c. 100 m thick) of probable earliest cambrian age. lithostratigraphic revision of this succession is presented here. the portfjeld formation throughout north greenland is elevated to the rank of group, and two new formations are defined in this group in southern peary land and immediately adjacent areas: the ediacaran carbonate-dominated unit is referred to the slusen formation, the upper mixed siliciclastic–carbonate unit to the glaciologelv formation. *correspondence: ji@geus.dk received: 07 may 2024 revised: 03 july 2024 accepted: 27 august 2024 published: 11 october 2024 keywords: neoproterozoic, slusen formation, glaciologelv formation, carbonate ramp, karst abbreviations: bace: basal cambrian excursion geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: karen dybkjær (geus, denmark) reviewed by: paul smith (oxford university museum of natural history, uk), keith dewing (natural resources canada, canada) funding: see page 18 competing interests: see page 18 additional files: none 1. introduction the basal sedimentary strata of the franklinian basin in southern peary land and adjacent areas of north greenland (figs 1, 2) were defined formally as the portfjeld formation by jepsen (1971; see also o’connor 1979; higgins et al. 1991a; ineson & peel 1997). this formation comprises a heterogeneous succession up to 300 m thick that is dominated by dolostones but includes prominent siliciclastic units and minor limestone intervals. strata referred to this formation are also recognised in north-eastern peary land and northern parts of central and western north greenland (christie & ineson 1979; peel 1982; higgins & soper 1985; surlyk & ineson 1987; higgins et al. 1991a; ineson & peel 1997), and inferred correlatives (referred to the ella bay formation) crop out on ellesmere island, nunavut, arctic canada (long 1989; dewing et al. 2004). the portfjeld formation was previously ascribed a probable early cambrian age, based on limited fossil evidence (peel 1988) and correlation to canada where lower cambrian fossils were reported by long (1989) from the kennedy channel formation that is overlain by the ella bay formation on judge daly peninsula. subsequent detailed mapping of ellesmere island demonstrated that the fossiliferous strata were emplaced tectonically from the ellesmere group, which overlies the ella bay formation; a probable neoproterozoic age was proposed for the ella bay formation and potentially for the correlative portfjeld formation in north greenland (dewing et al. 2004; peel & skovsted 2021). https://doi.org/10.34194/geusb.v57.8375 https://orcid.org/0000-0003-0017-3705 https://orcid.org/0000-0002-1774-7931 https://orcid.org/0000-0002-1434-7131 https://orcid.org/0000-0002-0040-6778 https://orcid.org/0000-0001-7385-4048 https://orcid.org/0000-0001-8062-0573 mailto:ji@geus.dk https://creativecommons.org/licenses/by/4.0/deed.ast ineson et al. 2024: geus bulletin 57. 8375. https://doi.org/10.34194/geusb.v57.8375 2 of 20 geusbulletin.org fig. 1 geological maps. a: the southern and northern outcrop belts of the portfjeld group across north greenland; inset shows the location of the study area and inglefield land (i) in northernmost greenland. fl: freuchen land. kb: kap bernhard. nf: nordenskiöld fjord. vgl: valdemar glückstadt land. red box indicates type area (fig 1b). b: the type area of the portfjeld group in southern peary land and heilprin land; modified from the digital greenland geological map at a scale of 1:500 000 (kokfelt et al. 2013). the locations (1–3) of the type and reference sections of the new formations (slusen formation, glaciologelv formation) are indicated: (1): type slusen formation, reference glaciologelv formation. (2): type glaciologelv formation, reference slusen formation. (3): reference glaciologelv formation. jbf: jørgen brønlund fjord. nmss: nedre midsommersø. ømss: øvre midsommersø. ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ 40°w 20°w 82°n 200 km ▲ ▲ ▲ permanent ice carboniferous – lower cenozoic (wandel sea basin) shelf trough northern outcrop belt: skagen group, portfjeld group southern outcrop belt: portfjeld group (slusen fm, glaciologelv fm) caledonian orogenic belt archaean middle–upper proterozoic fault caledonian frontal thrusts lower palaeozoic quaternary lake / fjord river cambro-ordovician buen formation (lower cambrian) morænesø formation (neoproterozoic) inuiteq sø formation (palaeo–mesoproterozoic) portfjeld group (slusen formation, glaciologelv formation) (ediacaran/lowermost cambrian) g.b. schley fjord independence fjord vgl kb j.c. christensen land mylius-erichsen land c al ed on ia n fo re la nd an d pa ra ut oc ht ho n peary land greenland ice sheet wandel dal heilprin land greenland i wulff land nf fl a b 82°15´n wandel dal ømss nmss slusen rundfjeld chr. erichsen iskappe storm iskappe glaciologelv jbf 3 82°n 10 km 37°w 35°w 33°w 2 1 https://doi.org/10.34194/geusb.v57.8375 http://www.geusbulletin.org/ ineson et al. 2024: geus bulletin 57. 8375. https://doi.org/10.34194/geusb.v57.8375 3 of 20 geusbulletin.org detailed sedimentological and stratigraphic investigation of the portfjeld formation in southern peary land in 2006 revealed the composite nature of the formation in this area, a prominent, regionally persistent karstic surface separating a lower carbonate-dominated unit (150–200 m thick) from an upper mixed siliciclastic–carbonate unit (c. 100 m thick; fig. 3a, b). preliminary isotopic data suggested a neoproterozoic age for the lower unit (ineson et al. 2007). recent biostratigraphic re-analysis of the lower carbonates revealed a rare but well-preserved micropalaeontological assemblage that includes acritarchs, cyanobacteria, and  putative eggs and embryos; in association with the δ13c isotope data, this microbiota confirms a late ediacaran (latest neoproterozoic) age (willman et al. 2020; rugen et al. 2022). although unfossiliferous, the upper unit of the portfjeld formation shows an isotopic signal that is compatible with an earliest cambrian age (willman et al. 2020). given this clear confirmation of the composite nature of the portfjeld formation in southern peary land, it is considered inadvisable to retain the definition of jepsen (1971) both on lithostratigraphic grounds and with a view to a more nuanced understanding of the early evolution of the franklinian basin. the aims of this paper, therefore, are first to elevate the portfjeld formation, as recognised widely in north greenland, to the rank of group, and second to define and describe two new formations in the portfjeld group in southern peary land and immediately adjacent areas (fig. 2). the detailed stratigraphy of these early franklinian basin deposits is of wider significance, however, both in charting patterns of early syn-rift subsidence and subsequent crustal break-up and as they host two palaeobiologically important lagerstätten – the recently reported ediacaran biota of the portfjeld group (willman et al. 2020) and the well-documented lower cambrian sirius passet biota in the overlying buen formation (peel 2010; ineson & peel 2011; harper et al. 2019). 2. regional setting northern greenland formed part of a major east– west-trending extensional basin, the franklinian basin, from the latest proterozoic to the devonian. this basin extended almost 2000 km across northern greenland and the canadian arctic islands; the two areas can be correlated closely at many stratigraphic levels (trettin 1989; higgins et al. 1991a, b; trettin et al. 1991; faehnrich et al. 2023). two discrete elements characterised the basin – a predominantly marine shelf to the south underlain by precambrian cratonic basement, passing north into a deep-water basin. the boundary between these two regimes was controlled by major east–west structural lineaments that were variably active during the history of the basin; the boundary shifted southwards with time, following the successive activation of these structures (surlyk et al. 1980; higgins et al. 1991a; surlyk 1991). the basin fill has a total thickness approaching 8 km; the shelf succession is carbonate-dominated, but an important siliciclastic formation is represented in the lower cambrian. deep-water basinal facies are largely siliciclastic with discrete intervals of redeposited this studyjepsen (1971) aftenstjernesø fm (brønlund fjord gp) buen fmbuen fm lo w er c am br ia n lo w er c am br ia n c am br ia n or eo ca m br ia n c (? ) ed ia ca ra n eo ca m b. pc m –p p ro t n glaciologelv fm po rt  e ld g ro up port eld fm slusen fm morænesø fm morænesø fm hfg inuiteq sø fm (independence fjord group) inuiteq sø fm brønlund fjord dolomite 20 19 18 10 –1 7 st ag e 4 st ag es 3 , 4 fig. 2 lithostratigraphic scheme of the precambrian – lower cambrian of southern peary land showing the relationship of the revisions presented here to the original definition of jepsen (1971); the numbers (10–20) refer to the lithological units in the type section described by jepsen (1971). c(?): inferred lowermost cambrian. eocamb.: eocambrian. hfg: hagen fjord group. m–p prot.: meso – palaeoproterozoic. n: neoproterozoic. pc: precambrian. fm: formation. gp: group. https://doi.org/10.34194/geusb.v57.8375 http://www.geusbulletin.org/ ineson et al. 2024: geus bulletin 57. 8375. https://doi.org/10.34194/geusb.v57.8375 4 of 20 geusbulletin.org b a buen fm buen fm uppermost slusen fm glaciologelv fm glaciologelv fm slusen formation proterozoic ssts/intrusions karst surface karst surface erosional unconformity, local regolith basal sandstones, glaciologelv fm karstic surface, top-slusen fm irregular karstic surface, top-glaciologelv fm 50 m 50 m fig. 3 general views of the portfjeld group in southern peary land. a: the portfjeld group (slusen and glaciologelv formations) on the northern side of rundfjeld, just south of the eastern end of nedre midsommersø, illustrating in particular the well-stratified, banded and cyclic nature of the slusen formation; the prominent dark unit in the lower slusen formation forms an important marker throughout the southern outcrop belt. b: view of the upper portfjeld group on the south side of øvre midsommersø showing the two reddened karstic surfaces bounding the glaciologelv formation. the internal architecture of the orange fluvial sandstones at the base of the buen formation records episodic syndepositional subsidence of the karstic depressions (dolines) at the top of the portfjeld group. https://doi.org/10.34194/geusb.v57.8375 http://www.geusbulletin.org/ ineson et al. 2024: geus bulletin 57. 8375. https://doi.org/10.34194/geusb.v57.8375 5 of 20 geusbulletin.org carbonate, shed from the coeval southern shelf regime. the franklinian basin was deformed and uplifted during the ellesmerian orogeny in the late devonian – early carboniferous. pervasive deformation and metamorphism were restricted to the deep-water basinal deposits in the north, though adjacent outermost shelf strata were involved in thin-skinned tectonics (fig. 1; soper & higgins 1990). the shelf deposits of southern peary land, under focus here, are undeformed, dipping gently towards the north. the extensional nature of the franklinian basin, evolving into a passive margin setting in the cambrian, is well documented (higgins et al. 1991a; surlyk 1991; hopper & ineson 2021), but the early syn-rift history and the timing of continental break-up remain poorly understood. the portfjeld group (as redefined here) is a key element in the early basin record, particularly as it is capped by a regional unconformity that is inferred to represent break-up and the onset of seafloor spreading (higgins et al. 1991a; surlyk 1991; hopper & ineson 2021). 3. lithostratigraphy the lithostratigraphy of the portfjeld formation of jepsen (1971) is formally revised here: the portfjeld formation is raised in rank to a group, and two new formations (slusen formation, glaciologelv formation) are defined in the southern outcrop belt. 3.1 lithostratigraphic framework of north greenland 3.1.1 portfjeld group revised unit history. defined and mapped as a formation in southern peary land by jepsen (1971) and recognised more widely in north greenland by christie & ineson (1979), o’connor (1979), peel (1982) and higgins & soper (1985). type area. wandel dal, southern peary land (fig. 1). thickness. in the type area of southern peary land, the group is 200–300 m thick, thinning south-eastwards to 50–70 m in eastern j.c. christensen land. in the northern outcrop belt, the thickness of the group ranges from 300 to 700 m, thickening northward to the northern limit of the group. lithology. the portfjeld group is a pale weathering, dolostone-dominated succession with subordinate sandrich siliciclastic units and minor limestones. in southern peary land, the group is well-stratified and banded; a prominent dark stripe dominates exposures along wandel dal (fig. 3a). varied microbial carbonates typify the type area, including thrombolitic bioherms and biostromes and mesoscopic domal stromatolites. cyclic grainstone-rich facies characterise certain levels, displaying intraclastic and ooidal fabrics; giant ooids occur at some levels. the grainstones display planar, trough and hummocky cross-stratification. siliciclastic sandstones occur locally at the base of the group and about 200 m above the base, within and above a prominent brecciated karstic surface. detailed descriptions and illustrations of the stratigraphic succession of the portfjeld group in the type area are provided here under the component formations. in the northern outcrop belt, the portfjeld group is characterised by a thick, uniform succession of pale grey dolostones displaying relict cross-stratified grainstone fabrics. subordinate microbial carbonates form discrete biostromal units that typically exhibit columnar stromatolites, although complex domal stromatolite patch reefs occur locally near the northern shelf margin, interdigitating with cross-bedded grainstones (ineson & peel 2011). fauna, flora and age. although essentially unfossiliferous and lacking an ichnofauna, a rare microbiota has been recovered from the lower portfjeld group in southern peary land; this includes acritarchs, cyanobacteria, putative eggs, problematic tubular fossils, embryos and red algal thalli (willman et al. 2020; see slusen formation, below). furthermore, the lower portfjeld group in this area exhibits a marked negative carbon isotope excursion (–12‰; willman et al. 2020; rugen et al. 2022). the palaeontological and isotopic data together indicate a general late ediacaran (neoproterozoic) age for this level of the group in southern peary land. fossils have not been recovered from the upper portfjeld group in southern peary land, nor from the entire group in the northern outcrop belt. the group is overlain unconformably by the buen formation, which has yielded fossils of early cambrian age (stages 3 and 4; peel & willman 2018), and the portfjeld group regionally is thus assigned a latest neoproterozoic – earliest cambrian age. depositional environment. the portfjeld group records deposition on a carbonate ramp and shallow-water carbonate platform that bordered the embryonic deep-water part of the franklinian basin (higgins et al. 1991a). details of the depositional facies and setting in the type area of southern peary land are provided below under the slusen and glaciologelv formations. in the northern outcrop belt, the portfjeld group is dominated by shallow-water, high energy platform facies that accumulated https://doi.org/10.34194/geusb.v57.8375 http://www.geusbulletin.org/ ineson et al. 2024: geus bulletin 57. 8375. https://doi.org/10.34194/geusb.v57.8375 6 of 20 geusbulletin.org in an outboard, marginal setting at the transition to the deep-water basin (surlyk & ineson 1987; higgins et al. 1991a; ineson & peel 1997, 2011). boundaries. in southern peary land, and to the southeast (heilprin land, j.c. christensen land), the portfjeld group unconformably overlies a range of precambrian units – the palaeoproterozoic–mesoproterozoic independence fjord group or the neoproterozoic morænsesø formation in southern peary land and northern heilprin land, the neoproterozoic hagen fjord group in south-east heilprin land and northern j.c. christensen land. the unconformity is characteristically flat but locally the basal strata onlap an irregular weathered regolith (see slusen formation, below). at the eastern end of the northern outcrop belt, the portfjeld group unconformably overlies precambrian sedimentary strata and volcanics, locally with significant relief (christie & ineson 1979). westwards, from north-west peary land to northern wulff land, the portfjeld group overlies with apparent conformity, the skagen group (higgins & soper 1985; surlyk & ineson 1987; higgins et al. 1991a). the portfjeld group is everywhere overlain unconformably by the lower cambrian buen formation (fig. 3). in southern peary land, along wandel dal (figs 1, 3), the unconformity is planar at a large scale, but highly irregular in detail, ranging from a mesoscale sutured surface to a series of isolated steep-sided depressions (dolines) up to several hundreds of metres across and tens of metres deep, filled with basal fluvial sandstones of the buen formation (fig. 3b). this karstic imprint at the unconformity is widespread, being also in evidence in north-west peary land (ineson & peel 2011; smith & moseley 2022) and central and western north greenland (davis & higgins 1987). at kap bernhard, north-east j.c. christensen land, in contrast, where the unconformity represents a more significant erosion surface (the upper portfjeld group being absent), the unconformity surface is flat and planar (see slusen formation, below). distribution. the portfjeld group crops out in two broad east–west belts (fig. 1a). the southern outcrop belt extends from the head of nordenskiöld fjord, south of freuchen land, eastwards to eastern valdemar glückstadt land (fig. 1a). excellent, continuous exposures occur on both sides of the valley of wandel dal in southern peary land. south-west of nordenskiöld fjord, the portfjeld group is not recognised in situ, but bedded slabs derived from the group form a prominent component of a megabreccia sheet up to 270 m thick that is well-exposed in southern wulff land (hurst & peel 1979; peel & wright 1985; surlyk & ineson 1987). the portfjeld group is absent in inglefield land in north-west greenland where the dallas bugt formation (equivalent to the buen formation) rests directly on precambrian rocks. the northern outcrop belt is discontinuous but is traceable from north-east peary land to northern wulff land within folded and/or faulted terranes, marginal to the north greenland fold belt. the portfjeld group crops out in faulted inliers in the east and in the cores of major anticlines in central and western north greenland. subdivision. in southern peary land and immediately adjacent areas to the south-east (heilprin land, j.c. chriestensen land, valdemar glückstadt land), the portfjeld group is herein formally subdivided into two formations, the slusen and glaciologelv formations. the portfjeld group in the northern outcrop belt is presently undifferentiated; detailed correlation with the southern area and potential formal subdivision necessitate future stratigraphical and geochemical investigations. 3.2 lithostratigraphy of the portfjeld group in southern peary land and adjacent areas 3.2.1 slusen formation new formation history. previously forming the lower two-thirds of the portfjeld formation in southern peary land as defined by jepsen (1971); the new slusen formation corresponds to units 10–17 of jepsen (1971; fig. 2). name. after slusen [the sluice], the short river that connects the two large elongate lakes øvre midsommersø and nedre midsommersø (figs 1b, 4). type section. cliff exposures on the north side of the river entering the head of øvre midsommersø (figs 1b [locality 1], 5, 6; 82°14´n, 36°06´w). reference section. 1 km north of the small peninsula in the middle of the northern shore of nedre midsommersø; the section broadly follows the gully paralleling an nnw–sse-trending dyke (figs 1b [locality 2], 7, 8; 82°15´n, 33°25´w). thickness. the slusen formation is 192 m thick in the type section (fig. 6). although semi-continuous exposures eastward along wandel dal indicate a broadly consistent thickness in this west–east profile, the slusen formation is only about 167 m thick in the reference section (fig. 8) due primarily to onlap onto a minor topographic high at the basal unconformity (see fig. 7). the formation is about 50–70 m thick at the northern end of the kap bernhard cliff section, eastern j.c. christensen land (o’connor 1979 and unpublished field notes and sections). https://doi.org/10.34194/geusb.v57.8375 http://www.geusbulletin.org/ ineson et al. 2024: geus bulletin 57. 8375. https://doi.org/10.34194/geusb.v57.8375 7 of 20 geusbulletin.org lithology. the slusen formation is a carbonate-dominated succession composed mainly of dolostones with subordinate limestones; siliciclastic facies occur locally at the base and top of the formation (figs 6, 8). the carbonates show a range of weathering colours, from pale silvery grey through fawn – khaki and mid grey – brown to dark grey or black. these commonly alternate cyclically on various scales to produce a strikingly banded or striped outcrop, which is readily traced across southern peary land (figs 3a, 5b, 7). the uppermost, karstified dolostones weather a rusty red-brown colour (fig. 3b). although mainly dolomitised, relict structures (often spectacularly preserved) record a spectrum of carbonate facies from cherty laminated mudstones through wackestones and ooid–intraclast packstones to ooid grainstones; the packstone–grainstone facies often show trough cross-bedding or hummocky cross-stratification. microbial fabrics are common, ranging from crinkly lamination and micro/mesoscale columnar and hemispherical domal stromatolites to thrombolitic bioherms and biostromes. a thin siliciclastic unit, up to 5 m thick, occurs locally at the base of the formation. it comprises mediumto coarse-grained sandstone, sometimes pebbly or glauconitic, showing planar cross-bedding in 5–20 cm sets, locally forming herringbone couplets, that indicate dominant northward directed palaeocurrents (kirkland et al. 2009). north of nedre midsommersø, the basal unconformity is markedly irregular (fig. 7), and the relict regolithic topography is onlapped by coarse pebble – cobble breccio-conglomerates that wedge out laterally within only a few metres; the pebble and cobble clasts are composed of quartzite derived locally from the underlying inuiteq sø formation (independence fjord group). the succeeding dolostone-dominated succession comprises two discrete sedimentary packets that often form separate topographic steps along the north side of the midsommer lakes. the lower packet (c. 120 m thick) begins with a 25–30 m thick succession of pale to mid grey weathering dolostones showing relict grainstone fabrics (ooid/peloid/intraclast) and trough cross-bedding (fig. 9a), with thin microbial (stromatolitic) beds in places. this basal unit darkens upwards, passing up into an interval (25–40 m thick) of stacked coarsening upward cycles, typically 2–5 m thick, comprising dark laminated dolostones (mudstone/wackestone) grading up into swaley and hummocky cross-stratified packestones and grainstones, locally capped by a thin stromatolitic bed (figs 6, 8). this cyclic succession passes up into a distinctive dark grey – black dolostone unit (typically 25–35 m thick) that is an important marker throughout the outcrop of the slusen formation (figs 3, 5–8, 9b; o’connor 1979); the fossil biota recorded by willman et al. (2020) was derived from samples collected near the base of this unit. although laminated, cherty carbonate mudstone fabrics dominate (fig. 9c), this dark dolostone unit also includes thin, often lenticular thrombolitic microbialites and isolated wackestone/packstone beds showing hummocky cross-stratification. the lower sedimentary packet of the slusen formation is capped by a thick (7–10 m), poorly stratified unit of thrombolitic dolostone that forms a prominent cliff in most sections (figs 3a, 5b). though a laterally persistent biostromal body in eastern exposures (fig. 3a), it shows a more complex fig. 4 slusen [the sluice], the short river linking the two lakes, øvre midsommersø in the distance and nedre midsommersø in the foreground (viewed towards the west). the portfjeld group is particularly well-exposed in the steep cliffs on both sides of øvre midsommersø. https://doi.org/10.34194/geusb.v57.8375 http://www.geusbulletin.org/ ineson et al. 2024: geus bulletin 57. 8375. https://doi.org/10.34194/geusb.v57.8375 8 of 20 geusbulletin.org b a fig. 5 slusen formation. a: view of the lower slusen formation at the type section above the western shores of the øvre midsommersø lake. stacked pale weathering grainstone-dominated cycles at the base of the formation are succeeded by darker mid-ramp and outer ramp facies (above encircled figure); a microfauna with late ediacaran affinities was obtained from samples taken laterally from the arrowed horizon. b: the slusen formation about a kilometre west of the type section showing the development of thrombolitic microbial mounds at a discrete horizon (arrowed) in the middle of the formation. https://doi.org/10.34194/geusb.v57.8375 http://www.geusbulletin.org/ ineson et al. 2024: geus bulletin 57. 8375. https://doi.org/10.34194/geusb.v57.8375 9 of 20 geusbulletin.org current cross-lamination wave cross-lamination carbonate mud drapes parallel lamination irregular, crinkly microbial lamination thrombolitic microbialite ooids, giant ooids columnar stromatolite domal stromatolite structures hummocky cross-strati�cation swaley cross-strati�cation trough cross-strati�cation (small-/medium-scale) trough cross-strati�cation (large-scale) planar cross-strati�cation platy intraclasts ( at-pebble conglomerate) g g black chert glaucony sandstone pebbly sandstone limestone intrusive sandstone dolostone dolomitic lithology mudstone/marlstone karstic breccia (dolostone/sandstone) white/light-coloured chert sl us en f or m at io n f ml m w g bd rudstone m c co gap g la ci ol og el v fo rm at io n m 270 260 250 240 230 220 210 200 190 180 170 160 150 140 g f ml m w g bd rudstone m c co mor. fm gap gap gap m sl us en f or m at io n 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 fig. 6 type section of the slusen formation and reference section of the glaciologelv formation, western øvre midsommersø (fig. 1b, locality 1). the slusen formation mounds illustrated in fig. 5b occur laterally to the type section at the 117 m level; due to access difficulties, the succeeding cyclic interval (123–136 m) was estimated from photographs. details of the primary cyclicity in the upper 50 m of the slusen formation are obscured by intense karst-related fracturing and brecciation. the glaciologelv formation reference section illustrates the fluvial sandstones at the base of the formation, draping the top-slusen formation karst surface. mor: morænesø. lithology/grain-size abbreviations: ml: marlstone. m: lime mudstone. w: wackestone. g: grainstone. bd: boundstone. rudstone subdivisions: f: fine. m: medium. c: coarse. co: cobble. the accompanying legend is applicable to figs 8 and 12. https://doi.org/10.34194/geusb.v57.8375 http://www.geusbulletin.org/ ineson et al. 2024: geus bulletin 57. 8375. https://doi.org/10.34194/geusb.v57.8375 10 of 20 geusbulletin.org architecture in the west, around øvre midsommersø, forming laterally impersistent mounds with depositional relief of up to 7 m that developed atop stratiform ooid grainstone bodies (fig. 9b, d). farther west, this unit loses its prominent, cliff-forming character. the upper sedimentary packet of the slusen formation (c. 75 m thick) is characterised by a lower, cyclic, mixed siliciclastic–carbonate succession that becomes more homogeneous and carbonate-dominated upwards and is capped by a rusty, locally irregular karstic surface (figs 3, 6, 8). the lower 25–30 m consists of well-developed cycles, 1.5–4 m thick, composed of limestones and/or dolostones interbedded with green-grey or purple marlstones (figs 3, 6, 8, 10a, b); individual cycles are rich in marlstone at the base and carbonate-dominated at the top. pale grey limestones are locally preserved in the lower 10–15 m, passing up or laterally across a diachronous front into brownish weathering dolostones (fig. 10a). the carbonate beds, typically 5–20 cm thick, comprise wackestones, packstones and intraclastic grainstones showing hummocky cross-stratification. bed bases are often defined by flat-pebble conglomerate divisions up to several tens of centimetres thick with scoured bases and megarippled undulating tops; complex, opposed imbrication fabrics are common. with decreasing siliciclastic content upwards, the succeeding c. 25–30 m thick dolostone succession comprises amalgamated hummocky and swaley cross-stratified cycles (figs 6, 8). these grade up into the uppermost slusen formation unit (25–30 m thick) that retains the cyclic pattern but is increasingly dominated up-section by microbial dolostones showing regular mesoscale, laterally linked hemispherical domes (0.5–2 m diameter, 20–40 cm synoptic relief; fig. 10c). where the karstic overprint is less penetrative, it is evident that this stromatolitic dolostone facies alone forms the upper 10–20 m of the formation. the karstic nature of the upper boundary of the slusen formation (see boundaries) becomes increasingly evident in the upper c. 20 m of the formation. dissolution cavities and vugs with early haematitic linings and drusy carbonate cements are common, and sub-vertical brecciated fracture planes coalesce to form irregular zones of clast-supported breccia (fig. 10d). sandstone matrix becomes increasingly common in the karstic breccias upwards. the uppermost levels of the formation are thus very variable laterally, passing from reddened and fractured yet in situ stromatolitic dolostones to broad zones (metres to tens of metres across) of sandy breccia, both clastand matrix-supported. fauna, flora and age. the slusen formation contains a variety of microscopic fossils, but no macroscopic fossils have been discovered (peel 1988; willman et al. buen fm glaciologelv fm slusen fm inuiteq sø fm (ifg) fig. 7 portfjeld group on the northern side of the nedre midsommersø lake, approximately 1 km west of the slusen formation reference section and the glaciologelv formation type section. at this locality, the basal unconformity is markedly irregular, the lower slusen formation onlapping prominent knolls of weathered proterozoic sandstone (small arrows). the upper glaciologelv formation is disrupted by a wide doline (large arrows) infilled with basal buen formation sandstone. ifg: independence fjord group. https://doi.org/10.34194/geusb.v57.8375 http://www.geusbulletin.org/ ineson et al. 2024: geus bulletin 57. 8375. https://doi.org/10.34194/geusb.v57.8375 11 of 20 geusbulletin.org f rudstone m c co ml m w g bdin d. f jo rd g ro up gap f ml m w g bd rudstone m c co gap gap gap gap sl us en f or m at io n gap m sl us en f or m at io n m g la ci ol og el v fo rm at io n bu en f or m at io n boundary estimated poor exposure 270 260 250 240 230 220 210 200 190 180 170 160 150 140 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 fig. 8 reference section of the slusen formation and type section of the glaciologelv formation, nedre midsommersø (fig. 1b, locality 2), exhibiting the cyclic nature of both the slusen formation and the carbonates of the glaciologelv formation. ind.: independence. lithology/grain-size abbreviations: ml: marlstone. m: lime mudstone. w: wackestone. g: grainstone. bd: boundstone. rudstone subdivisions: f: fine. m: medium. c: coarse. co: cobble. for legend, see fig. 6. https://doi.org/10.34194/geusb.v57.8375 http://www.geusbulletin.org/ ineson et al. 2024: geus bulletin 57. 8375. https://doi.org/10.34194/geusb.v57.8375 12 of 20 geusbulletin.org fig. 9 lower slusen formation, representative lithologies from the type area, øvre midsommersø. a: trough cross-bedded grainstones (dolostones); type section, 23 m above base. b: dark laminated dolostones and black cherts; type section, 83 m above base. c: slusen formation, c. 1 km west of the type section, showing the abrupt boundary (equivalent to 92.5 m in the type section) between the dark outer ramp dolostones and the overlying cyclic ooidal grainstones and capping thrombolitic mounds. the relief on the arrowed mound is about 5 m. d: giant ooids; grainstones immediately beneath the mounds illustrated in fig. 9c. pencil tip for scale. e: internal thrombolitic fabric typical of the microbial mounds illustrated in fig. 9c; photograph from a fallen slab about 4 km west of the type section. a c d e b 10 cm https://doi.org/10.34194/geusb.v57.8375 http://www.geusbulletin.org/ ineson et al. 2024: geus bulletin 57. 8375. https://doi.org/10.34194/geusb.v57.8375 13 of 20 geusbulletin.org 2020; peel & willman 2022; willman & peel 2022). the microscopic, phosphatised fossils were derived from dolostone near the base of the distinctive dark dolostone unit in the lower part of the formation (fig. 5a) and consist of a three-dimensionally preserved biota including acritarchs, cyanobacteria, red algal thalli, problematic tubular fossils and putative eggs and embryos. this well-preserved and diverse assemblage is similar in composition and preservation to the ediacaran age doushantuo formation lagerstätte in china (e.g. xiao et al. 1998, 2014; zhang et al. 1998). a negative carbon isotope excursion of –12‰ expressed across the reference section of the slusen formation currently supports a late ediacaran age assignment (willman et al. 2020; rugen et al. 2022). a precise correlation with the global neoproterozoic carbon isotope profile remains unclear, however, as the nature of the chemostratigraphic trend shares similarities with a dc b fig. 10 upper slusen formation, representative lithologies from the type area, øvre midsommersø. a: marlstone – limestone/dolostone cycles equivalent to the 123–140 m interval in the type section (fig. 6); photograph about 4 km west of the type section, the cycles are 2–3 m thick. the weathering change from pale grey (limestone) to orange (dolostone) at the top of the image (arrowed) marks the dolomite front. b: close-up of the marlstone – limestone/dolostone cyclic facies; the thin, lenticular carbonate beds (e.g. at the hammer) commonly display hummocky cross-stratification. c: mesoscale domal stromatolites characterise the upper levels of the slusen formation (172–194 m in the type section (fig. 6) and 141–167 m in the reference section (fig. 8)); photograph from c. 1 km east of the type section. d: karstic breccia, uppermost slusen formation, type section. https://doi.org/10.34194/geusb.v57.8375 http://www.geusbulletin.org/ ineson et al. 2024: geus bulletin 57. 8375. https://doi.org/10.34194/geusb.v57.8375 14 of 20 geusbulletin.org both the global shuram excursion of late ediacaran age and the basal cambrian excursion (bace; rugen et al. 2022). the combined palaeontological and isotopic evidence indicates a general late ediacaran, neoproterozoic age for the slusen formation. depositional environment. the facies represented in the slusen formation testify to a storm-dominated carbonate ramp setting. the shallow-water inner ramp was characterised by siliciclastic tidal sands, carbonate sand shoals and stromatolite tracts, the mid-ramp by carbonate storm sands, and the outer ramp by carbonate muds. microbial bioherms and biostromes developed during a major ramp drowning event. boundaries. in southern peary land and northern heilprin land, the slusen formation rests unconformably on sandstones of the inuiteq sø formation, part of the palaeoproterozoic–mesoproterozoic independence fjord group (fig. 11a), or locally on erosionally truncated outliers of the morænesø formation (neoproterozoic tillites and carbonates of inferred marinoan affinity; kirkland et al. 2009); for details of proterozoic stratigraphy, see sønderholm & jepsen (1991). being observed in cliff sections, the basal unconformity appears planar on a large scale but locally shows significant relief. on the northern shores of nedre midsommersø, for example, relief of up to 30 m is evident, the lower dolostones of the slusen formation onlapping knolls of deeply weathered regolithic sandstone of the independence fjord group (fig. 7). this positive relief is locally highly asymmetric, with steep, locally subvertical western and northern slopes, possibly representing degraded fault scarps. south-east of the type area, the slusen formation unconformably overlies shelf strata on the western margin of the hekla sund basin, a neoproterozoic extensional basin of iapetus affinities (clemmensen & jepsen 1992; sønderholm et al. 2008). in south-eastern heilprin land (fig. 1), the slusen formation unconformably overlies the catalinafjeld formation (hagen fjord group), whereas at kap bernhard, north-east j.c. christensen land, the slusen formation rests unconformably on dolostones of the fyn sø formation, the uppermost formation of the hagen fjord group (fig. 11b). the unconformity surface at kap bernhard is planar although exposures at the northern end of the cliff-line reveal normal faults that displace strata of the fyn sø formation yet are truncated at the base slusen formation unconformity. in southern peary land, and probably west to the head of nordenskiöld fjord, south of freuchen land (fig. 1), the upper boundary of the slusen formation is a distinctive, reddened karstic unconformity, overlain by cross-bedded coarse-grained pebbly sandstones or poorly exposed fine-grained siliciclastic sediments at the base of the glaciologelv formation, as defined here (figs 3b, 6, 8). although sub-planar and occurring at a consistent stratigraphic level, this unconformity is locally irregular in detail with prominent mounds of relict dolostone projecting up to 5 m above depressions floored by sandy karstic breccia. at kap bernhard, north-east j.c. christensen land, the glaciologelv formation is absent due to regional south-eastward truncation at the base of the buen formation (upper lower cambrian); the slusen formation is thus overlain unconformably by the buen formation at this locality (fig. 11b). distribution. the slusen formation is recognised from the head of nordenskiöld fjord, south of freuchen land, eastwards along wandel dal in southern peary land to easternmost valdemar glückstadt land. accessible, well-exposed sections are particularly common along the northern shores of the midsommer lakes in wandel dal. exposure in the eastern outcrops, in j.c. christensen land and valdemar glückstadt land, is generally poor, with the exception of the precipitous cliffs at kap bernhard (fig. 11b). 3.2.2 glaciologelv formation new formation history. the succession referred here to the new glaciologelv formation previously formed the upper third of the portfjeld formation in southern peary land as defined by jepsen (1971); the glaciologelv formation corresponds to units 18–20 of jepsen (1971; fig. 2). type section. about 1.5 km north of the small peninsula in the middle of the northern shore of nedre midsommersø (figs 1b [locality 2], 7, 8; 82°15´n, 33°25´w); the section is at the head of the gully where the cross-bedded, grainstone-dominated dolostones that typify the formation are well-exposed. reference sections. accessible, complete sections through the glaciologelv formation are scarce, and the type section at nedre midsommersø is supplemented by reference sections immediately north of the western end of øvre midsommersø (figs 1b [locality 1], 6; 82°14´n, 36°06´w), and along the southern bank of glaciologelv, a small river south of the head of jørgen brønlund fjord (figs 1b [locality 3], 12; 82°10´n, 31°18´w). the former illustrates details of the basal siliciclastic interval and the latter the uppermost c. 20 m of the formation, intervals that are incomplete or covered at the type section. https://doi.org/10.34194/geusb.v57.8375 http://www.geusbulletin.org/ ineson et al. 2024: geus bulletin 57. 8375. https://doi.org/10.34194/geusb.v57.8375 15 of 20 geusbulletin.org thickness. accessible sections are generally incomplete (see figs 6, 8) but based on lateral correlation between type and reference sections, the glaciologelv formation is estimated to be about 100 m thick in southern peary land. the thickness probably varies by at least 20–30 m laterally, however, due to the clear irregularity of the karstic surfaces at the base and top of the formation, as observed in inaccessible cliff sections (fig. 3). buen fm glaciologelv fm slusen fm buen fm slusen fm fyn sø fm inuiteq sø fm b a fig. 11 a: portfjeld group at northern rundfjeld, south of nedre midsommersø, showing the basal unconformity overlying independence fjord group quartzites and sills (dark, right); central conglomeratic wedge in the inuiteq sø formation potentially related to pre-portfjeld group faulting (see fig. 7). b: the slusen formation at kap bernhard, ne j.c. christensen land (fig. 1a), unconformably overlies the karstified(?) fyn sø formation (neoproterozoic hagen fjord group) and is overlain at a sharp flat unconformity by the buen formation. the glaciologelv formation is absent in this eastern area. https://doi.org/10.34194/geusb.v57.8375 http://www.geusbulletin.org/ ineson et al. 2024: geus bulletin 57. 8375. https://doi.org/10.34194/geusb.v57.8375 16 of 20 geusbulletin.org lithology. the glaciologelv formation is composed of three discrete elements that are recognised in all studied sections, albeit variably exposed and developed. the basal siliciclastic unit, 15–30 m thick, forms a recessive weathering ledge above the karstified top of the slusen formation (figs 3, 7). it comprises a laterally variable alternation of prominent, rusty weathering lenticular sandstone bodies and more recessive finer-grained deposits. the sand bodies are up to 8 m thick and comprise pebbly coarseto very coarse-grained sandstones showing large-scale trough cross-bedding (fig.  13a), grading up to trough and planar cross-bedded, mediumto fine-grained sandstone and current-rippled fine-grained sandstones and heteroliths; the sedimentary structures indicate sediment transport towards the north (kirkland et al. 2009). the middle and dominant unit of the formation (c. 50 m thick) comprises a uniform succession of mid grey-brown weathering dolostones, often reddening upwards, that show relict grainstone fabrics (coarse sand grade) and are typically trough cross-bedded (figs 8, 13b). finer-grained, hummocky cross-stratified intervals, c. 0.5 m thick, alternate with trough cross-bedded dolostone packets up to 6 m thick in the lower 30 m; the former intervals thin upwards and above 30 m, the succession is composed solely of the coarse, cross-bedded grainstone facies (fig. 8). a distinctive c. 2 m thick columnar stromatolite bed marks this shift from cyclic to amalgamated grainstones (figs 8, 13c). the upper boundary of this middle unit is a sharp, irregular scalloped surficial karst surface (karren), well exposed in the glaciologelv section, south of brønlund fjord (figs 12, 14a). the upper unit of the glaciologelv formation is a c. 20 m thick succession of cross-bedded siliciclastic sandstones, often with carbonate intraclasts and a dolomite cement, and sandy dolostones. the unit is covered in the type section and is best observed at glaciologelv, south of jørgen brønlund fjord (figs 1b, 12). sandwiched between the surficial exposure surface at the base (fig.  14a) and the highly irregular karstic profile at the top (the glaciologelv formation – buen formation unconformity, see boundaries), the depositional fabric of this unit is often obscure. in addition to karst-related collapse, the succession also shows evidence of in situ brecciation and extensive sand fluidisation and injection (figs 12, 14b). the primary sedimentary lithologies are cross-bedded mediumto coarse-grained, locally coarse to very coarse pebbly sandstones interbedded with sandy dolostones (grainstones). trough and less commonly planar cross-sets, locally in herringbone couplets, are typically 5–20 cm thick and indicate bimodal but dominantly eastward-directed palaeocurrents (kirkland et al. 2009); hummocky cross-stratification is observed locally. fauna/flora and age. fossils are not known from the glaciologelv formation. its age is constrained stratigraphically by the ediacaran biota of the underlying slusen formation (willman et al. 2020) and by lower cambrian (cambrian series 2, stage 3–4) fossils from the upper member of the overlying buen formation (vidal & peel 1993; peel & willman 2018; wallet et al. 2021). trace fossils indicative of a lower cambrian age are common in sandstones and siltstones in the upper part of the g la ci ol og el v fo rm at io n (u pp er ) bu en f m m ml m carbonate pebblesand siliciclastic w g fmud m mfvcvf c c 20 15 10 0 5 fig. 12 reference section of the glaciologelv formation on the southern bank of the glaciologelv river, south of jørgen brønlund fjord (fig. 1b, locality 3). this section illustrates the uppermost levels of the formation (generally poorly exposed or inaccessible in the midsommersø sections) and details of the irregular karstified unconformity with the buen formation. lithology/grain-size abbreviations: carbonate: ml: marlstone. m: lime mudstone. w: wackestone. g: grainstone. siliciclastic: vf: very fine. f: fine. m: medium. c: coarse. vc: very coarse. for legend, see fig. 6. https://doi.org/10.34194/geusb.v57.8375 http://www.geusbulletin.org/ ineson et al. 2024: geus bulletin 57. 8375. https://doi.org/10.34194/geusb.v57.8375 17 of 20 geusbulletin.org lower member of the buen formation (bryant & pickerill 1990). the stratigraphic expression of carbon isotopes at the type section of the glaciologelv formation has been suggested to best correlate with the global cambrian, fortunian record (rugen et al. 2022). the glaciologelv formation is thus considered to be of probable earliest cambrian age. depositional environment. the basal siliciclastic succession represents a fluvial setting (channels and overbank/interfluve), whereas the middle dolostones attest to high energy, shallow marine processes on a carbonate platform, probably shallowing from middle–upper shoreface to the intertidal zone and capped by an ephemeral exposure surface. the succeeding mixed siliciclastic–carbonate deposits are also shallow marine in origin, influenced both by storms and tides. boundaries. the glaciologelv formation is bounded by two karstic unconformities, often delineated by distinctive reddened, or locally greenish, horizons (figs 3, 11a, 14c). the base (see slusen formation) is a subplanar surface but irregular in detail with up to at least 5 metres of demonstrable relief where positive zones of fractured yet broadly in situ upper slusen formation dolostone rise above dolostone breccia with a sandstone matrix. siliciclastic sediments at the base of the glaciologelv formation drape this irregular surface. the upper karstic unconformity between the glaciologelv formation and the lower cambrian (cambrian series 2, stage 3–4) buen formation is strikingly irregular, ranging from a sharp, metre-scale sutured surface to deep depressions up to several hundreds of metres across and tens of metres deep filled with basal buen formation fluvial sandstones (figs 3b, 7, 14c). distribution. the glaciologelv formation is recognised from the head of nordenskiöld fjord in the west across southern peary land, particularly in the type area in the valley of wandel dal, to the northern shores of independence fjord in the east (fig. 1a). farther eastward, the glaciologelv formation is probably present, albeit thin (<20 m), in northern j.c. christensen land (o’connor 1979 and unpublished field notes) but is apparently truncated south-east of these exposures. at kap bernhard, eastern j.c. christensen land, the glaciologelv formation is absent, and the buen formation rests directly on the slusen formation (fig. 11b); a similar relationship is likely in eastern valdemar glückstadt land (o’connor 1979 and unpublished field notes). a c b fig. 13 lower–middle glaciologelv formation, representative lithologies. a: pebbly coarseto very coarse-grained sandstones at the base of the  glaciologelv formation showing large-scale trough cross-bedding; person for scale (right). øvre midsommersø reference section, 190– 198 m. b: trough cross-bedded grainstones (dolostones), type section 235 m; pencil (top right) for scale. c: columnar stromatolites, type section, 232 m. https://doi.org/10.34194/geusb.v57.8375 http://www.geusbulletin.org/ ineson et al. 2024: geus bulletin 57. 8375. https://doi.org/10.34194/geusb.v57.8375 18 of 20 geusbulletin.org acknowledgements we extend our thanks to jørgen a. bojesen koefoed, chris kirkland and lars stemmerik for assistance and companionship in the field, and to stine øckenholt and jette halskov for graphical inputs. funding for fieldwork in 2006 from the danish natural science research council is gratefully acknowledged. we also acknowledge the pioneer stratigraphic studies on this succession carried out by hans jepsen and bernard o’connor. this paper has benefitted greatly from the detailed and constructive comments provided by the two referees. additional information funding statement funding for fieldwork in 2006 was supported by the danish natural science research council (fnu jr. nr. 272-05-0293). author contributions jri: funding acquisition; conceptualisation; investigation; visualisation; writing – original draft; writing – review and editing. jsp: conceptualisation; investigation; writing – original draft; writing – review and editing. sw: investigation; writing – original draft. ejr: investigation; writing – original draft. ms: investigation. pf: investigation. competing interests the authors declare no competing interests. references bryant, i.d. & pickerill, r.k. 1990: lower cambrian trace fossils from the buen formation of central north greenland: preliminary observations. rapport grønlands geologiske undersøgelse 147, 44–62. https://doi.org/10.34194/rapggu.v147.8106  christie, r.l. & ineson, j.r. 1979: precambrian–silurian geology of the g.b. schley fjord region, eastern peary land, north greenland. rapport grønlands geologiske undersøgelse 88, 63–71. https://doi. org/10.34194/rapggu.v88.7557  a c b fig. 14 upper glaciologelv formation. a: surficial karstic surface (karren) at the boundary between the cross-bedded dolostones of the middle glaciologelv formation and the overlying dolomitic sandstones; 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https://doi.org/10.1002/spp2.1347 https://doi.org/10.1017/jpa.2022.43 https://doi.org/10.1016/j.precamres.2022.106781 https://doi.org/10.1016/j.precamres.2022.106781 https://doi.org/10.1038/35318 https://doi.org/10.1666/12-157r https://doi.org/10.1017/s0022336000059977 https://doi.org/10.1017/s0022336000059977 lithostratigraphy of the portfjeld group (ediacaran lowermost cambrian) of north greenland 1. introduction 2. regional setting 3. lithostratigraphy 3.1 lithostratigraphic framework of north greenland 3.1.1 portfjeld group 3.2 lithostratigraphy of the portfjeld group in southern peary land and adjacent areas 3.2.1 slusen formation 3.2.2 glaciologelv formation acknowledgements additional information funding statement author contributions competing interests references figures fig. 1 geological maps. a: the southern and northern outcrop belts of the portfjeld group across n fig. 2 lithostratigraphic scheme of the precambrian lower cambrian of southern peary land showing fig. 3 general views of the portfjeld group in southern peary land. a: the portfjeld group (slusen a fig. 4 slusen [the sluice], the short river linking the two lakes, øvre midsommersø in the distance fig. 5 slusen formation. a: view of the lower slusen formation at the type section above the western fig. 6 type section of the slusen formation and reference section of the glaciologelv formation, wes fig. 7 portfjeld group on the northern side of the nedre midsommersø lake, approximately 1 km west o fig. 8 reference section of the slusen formation and type section of the glaciologelv formation, ned fig. 9 lower slusen formation, representative lithologies from the type area, øvre midsommersø. a: t fig. 10 upper slusen formation, representative lithologies from the type area, øvre midsommersø. a: fig. 11 a: portfjeld group at northern rundfjeld, south of nedre midsommersø, showing the basal unco fig. 12 reference section of the glaciologelv formation on the southern bank of the glaciologelv riv fig. 13 lower-middle glaciologelv formation, representative lithologies. a: pebbly coarseto very c fig. 14 upper glaciologelv formation. a: surficial karstic surface (karren) at the boundary between geological survey of denmark and greenland bulletin 20, 2010, 83–86 83 the greenland ice sheet is one of the most significant contributors to the rising global sea level with a contribution of 0.5 mm per year (rignot & kanagaratnam 2006). evidence is emerging that rising temperatures of subsurface ocean currents play a vital role in the recent acceleration of large fast flowing glaciers such as jakobshavn isbræ in west greenland (holland et al. 2008) and helheimgletscher in south-east greenland (straneo et al. 2010). important questions are whether these incursions of warmer water are part of a recurrent phenomenon and indeed exactly how they influence the glaciers. the geocenter denmark project sedimice (linking sediments with ice-sheet response and glacier retreat in greenland) investigates past ice fluctuations in the helheimgletscher region in south-east greenland with regard to magnitude, possible causes and effects. one of the main tasks in this project is to analyse sedimentary deposits in the main fjord sermilik (fig. 1), which is influenced by the tidally affected helheimgletscher that has a short floating tongue. by combining sediment studies with modern climate studies we aim to extrapolate meteorological data back in time. in august 2009 the geological survey of denmark and greenland collected short sediment cores in sermilik near tasiilaq (fig. 1). to select core sites and to understand the sedimentary processes, we also acquired data on the bathymetry and conducted shallow seismic profiling. this paper presents some results of the seismic survey, preliminary sediment core data and bathymetrical data from the fjord. apart from a few isolated depth values, the bathymetry of sermilik was unknown before the 2009 survey. setting sermilik is about 80 km long and 7–13 km wide. the terrain around the fjord is alpine with elevations of 300–600 m near the coast and up to 1000 m inland. frequent glacial and geologically controlled fissure valleys dissect the area in a criss-cross pattern. the northern end of sermilik branches into three fjords with calving glaciers. the westernmost – helheimgletscher – is a fast flowing glacier and the third most prolific iceberg producer in greenland (rignot & kanagaratnam 2006). the climate of the region is low arctic and the weather conditions are influenced by lows moving north along the coast. the fjord is covered by sea ice from december to may. the hydrographic conditions in the fjord are influenced by a 10– 20 m thick layer of glacial water, underlain by 100–150 m of bathymetry, shallow seismic profiling and sediment coring in sermilik near helheimgletscher, south-east greenland camilla snowman andresen, niels nørgaard-pedersen, jørn bo jensen and birger larsen fig. 1. study area and bathymetrical data from sermilik based on data collected in 2009. depth data on the shelf south of the fjord mouth are from clausen (1998) and this survey. the positions of the seismic lines a, b and c are indicated by black lines and sediment core sample sites are shown. 37°0´w toqqulertivit imiat shelf trough tasiilaq sill er07 c b a er15 er13 er11 10 km bank 700 m 200 m helheimgletscher 38°w 66°n se rm ilik greenland 3–100 101–200 201–300 301–400 401–500 501–600 601–700 701–800 801–900 901–1000 1001–1100 sample site bathymetry depth (m) depotø © geus, 2010. geological survey of denmark and greenland bulletin 20, 83–86. open access: www.geus.dk/publications/bull 8484 polar water. below the polar water towards the bottom warmer atlantic water of subtropical origin with temperatures of 3.5–4°c is found (straneo et al. 2010). the inflow of warmer waters into sermilik takes place via deep troughs in the shelf. glacial history of the region a study from the toqqulertivit imiat valley (fig. 1) shows that a glacier flowed through this valley and most likely coalesced with a glacier flowing south in sermilik and out over the continental shelf during the last glacial maximum (roberts et al. 2008). exposure ages of 11.8–9.9 ka (kilo-annum, 103 years bp) from bedrock surfaces at high elevations (683–740 m a.s.l.) provide minimum ages for the last deglaciation (roberts et al. 2008). these ages from toqqulertivit imiat support the ‘maximum’ model of a large last glacial maximum ice sheet extending to the shelf break in southeast greenland (stein et al. 1996; kuijpers et al. 2003). evidence from the shelf south of sermilik indicates that the ice margin retreated to the inner shelf around 15.7–14.6 calibrated (cal.) ka (kuijpers et al. 2003). in the kangerlussuaq region farther north the ice-sheet margin retreated from the outer shelf around 15.5 cal. ka and reached the present outer coast around 13.6–10 cal. ka (jennings et al. 2006). this is in accordance with surface exposure ages from lower toqqulertivit imiat indicating that ice retreated to the mouth of sermilik between 11.1 and 9.7 ka (roberts et al. 2008). these data are further supported by a minimum age of 11 cal. ka for the formation of the local marine limit (at 69 m) and thereby local ice retreat near tasiilaq (long et al. 2008). methods we used the locally hired motor boats erik den røde and puite for the work. an innomar ses-2000 medium sub-bottom echo sounder from innomar tecnologies, rostock, germany, was used for bathymetrical and sub-bottom sediment profiling. this parametric device is designed for water depths down to about 2000 m and has the ability to resolve sediment layers a few decimetres thick and penetrate down to about 50 m below the sea floor. the transducer was mounted on a vertical steel tube attached to the side of the boat and a motion sensor was used to compensate for movements of the boat. additional bathymetrical data were recorded in the inner fjord from the echo sounder of erik den røde during the sediment coring cruise. comparisons of depth recordings obtained by the two methods showed that the results are compatible to within a few metres. sediment coring was performed with a rumohr lot corer with up to 1.5 m long core liners. bathymetrical data from sermilik the southern part of sermilik fjord is an up to 920 m deep and flat basin (fig. 1). the bathymetry of the fjord mouth can be described as terminating into a se-directed trough and a sw-directed trough separated by a broad bank with water depths of less than 200 m. the bathymetry of the swdirected trough is very uncertain. the shallowest part of the se-directed trough forms a c. 550 m deep sill between the deep fjord and the 800–900 m deep trough that stretches the entire shelf towards the irmiger sea. the deep fjord basin extends up to 40 km northward from the fjord entrance into the middle part of sermilik where several bathymetrical highs (400–550 m) narrow the connection to the northern part of the fjord. towards the northern part of sermilik, the basin floor rises steadily to a depth of about 600–650 m just south of depotø. the fjord bottom in the inner part is more irregular with channel systems more than 100 m wide and up to 20 m deep. during the survey we could not measure water depths in the inner east– west-trending fjord arm leading up to helheimgletscher due to semi-permanent sea ice extending tens of kilometres er15 er13 er11er07 5 cm fig. 2. selected examples of x-ray radiographs from core er15 (600 m water depth), er13 (660 m water depth), er07 (525 m water depth) and er11 (600 m water depth) that document different sedimentation regimes. note the dark sand layer in core er13 with a lower erosive boundary; this unit is interpreted to be a turbidite. core sites are shown in fig.1. 85 from the helheimgletscher calving front. however, according to the skipper of erik den røde (sigurdur petursson), water depths up to 800 m are found north-west of depotø. sediment cores altogether 19 cores with lengths ranging from 30 to 150 cm were retrieved during the sediment coring cruise. the full sediment core data (sedimentology, geochemistry and chronology) will be presented elsewhere. however, preliminary inspection of cores er07, er11, er13 and er15 documents the variable sediment regimes that characterise the fjord (figs 1, 2), and x-ray radiography of the cores show diamicton facies, laminated mud facies and sand layers with erosive lower boundaries. these lithofacies are similar to lithofacies described from sediment cores from kangerlussuaq (smith & andrews 2000) and scoresby sund (dowdeswell et al. 1994; ó cofaigh & dowdeswell 2001). for example, core er15 consists of laminated mud with variable content of pebbles, which is interpreted as ice-proximal glaciomarine sediments mainly deposited by suspension settling from turbid overflow plumes and turbidity currents and occasional iceberg rafting. in contrast, cores er07 and er11 are characterised by massive diamicton facies with abundant pebbles, which is interpreted as the result of iceberg rafting. core er13 has a unit of diamicton facies above a unit of laminated mud facies. this may reflect an environmental change from a long-lasting sea-ice cover in the fjord prohibiting iceberg passage to a period with increased passage and melting of icebergs. as also suggested by jennings & weiner (1996) variable inflow of atlantic water may influence the melting and traversing of icebergs. 210pb dating of the upper decimetres of the cores show sedimentation rates >1 cm/yr in er15, 0.4 cm/yr in er13 and 0.2 cm/yr in er11. the decreasing sedimentation rates with increasing distance to the present front of the calving glaciers reflect the decreasing influence from meltwater plume sedimentation. seismic profiles the seismic profile (fig. 3a) shows an outer flat, deep basin in sermilik with an upper 4–6 m thick, transparent seismic unit overlying a well-stratified section (>15 m thick) that is characterised by strong, continuous, parallel reflectors. there is no distinct boundary between the two seismic units, as weak, discontinuous, parallel reflectors characterise the uppermost c. 2 m of the stratified section. on vertically extremely exaggerated sections, wide lenticular units and stratigraphical onlap structures are visible in some parts of the lower seismic unit. to the north, the transparent upper unit disappears and well-stratified sediments, with some channel features, dominate the seismic profiles (fig. 3b). the profiles of the bathymetrical highs in the middle part of the fjord are dominated by overlapping diffraction hyperbolae. the inner, shallower part of the fjord is characterised by large channel and levee structures, and broad flank units showing well-stratified sediments in the north-western part of the seismic survey area (fig. 3c). formation of fjord-bottom sediment structures the sediment unit with a transparent pattern in the outer fjord basin indicates a uniform lithology that possibly originwnw sese nwnw sese nw se nw se nn ssn s c b a 1 km10 m 1 km5 m 1 km 5 m fig. 3. representative seismic sections that illustrate different structures in different parts of sermilik (see text for description and fig. 1 for location). 8686 nates from suspension sedimentation (meltwater plumes) and ice-rafting during the main part of the holocene (fig. 3a). the lower stratified section is interpreted as glaciomarine sediments consisting of turbidites and mass-transport deposits interbedded with suspension-deposited sediments. the lower unit was possibly deposited during the final deglaciation of the main part of sermilik at about 10 ka (cf. roberts et al. 2008). turbidite sedimentation typically creates very flat fjord basins with reflectors onlapping basin margins and structural highs. it is possible that the structural highs in the middle part of sermilik could serve as anchor points for the retreating glacier, causing a stagnation of the fjord glacier front during the last deglaciation. the inner basin with its apparent active channel and levee sedimentation and over-all fill geometry (fig. 3c) can be characterised as a progradational–aggradational wedge of sediments with possibly very high sedimentation rates from turbidites and mass flows, as well as plume sedimentation. it is an open question whether the channel systems are directly fed by the helheimgletscher source, or whether bedrock thresholds in the innermost fjord system prohibit bed transport of glaciogenic sediments. if a deeper sub-basin exists north of depotø, we can only explain the seismic signature and large channel-levee systems of the inner basin by a very advanced position of helheimgletscher to near depotø during the late holocene. hopefully, future exposure ages from the land terrain near depotø by our collaborators can show if the front of helheimgletscher had a standstill near depotø during the little ice age. in conclusion, the seismic survey has revealed a rather complex pattern of sub-bottom sediment structures (down to 50 m) in sermilik reflecting the early holocene retreat of helheimgletscher, probably followed by a holocene ice advance – perhaps during the little ice age. preliminary results from analyses of the sediment cores support the interpretation that the glaciomarine sediments in sermilik is related to settling from meltwater plumes and iceberg rafting. knowledge of the sediment depositional regime on the fjord bottom from seismic investigations is highly relevant for retrieval of sediment cores suitable for holocene palaeoclimatic reconstructions. we hope to collect more and longer sediment cores in the fjord in coming years. acknowledgements geocenter denmark is thanked for financial support. we would also like to thank our two skippers bendt josvassen of m/v puite and sigurdur petursson of m/v erik den røde. rineke gieles at the royal netherlands institute for sea research is thanked for x-ray radiography of sediment cores. references clausen, l. 1998: the southeast greenland glaciated margin: 3d stratal architecture of shelf and deep sea. geological society special publications (london) 129, 173–203. dowdeswell, j.a., whittington, r.j. & marienfeld, p. 1994: the origin of massive diamicton facies by iceberg rafting and scouring, scoresby sund, east greenland. sedimentology 41, 21–35. holland, d.m., thomas, r.h., de young, b., ribergaard, m. h. & lyberth, b. 2008: acceleration of jakobshavn isbræ triggered by warm subsurface ocean waters. nature geoscience 1, 659–664. jennings, a.e. & weiner, n.j. 1996: environmental change in eastern greenland during the last 1300 years: evidence from foraminifera and lithofacies in nansen fjord, 68°n. the holocene 6, 179–191. jennings, a.e., hald, m., smith, m. & andrews, j.t. 2006: freshwater forcing from the greenland ice sheet during the younger dryas: evidence from southeastern greenland shelf cores. quaternary science reviews 25, 282–298. kuijpers, a., troelstra, s.r., prins, m.a., linthout, k., akhmetzhanov, a., bouryak, s., bachmann, m.f., lassen, s., rasmussen, s. & jensen, j.b. 2003: late quaternary sedimentary processes and ocean circulation changes at the southeast greenland margin. marine geology 195, 109–129. long, a.j., roberts, d.h., simpson, m.j.r., dawson, s., milne, g.a. & huybrechts, p. 2008: late weichselian relative sea-level changes and ice sheet history in southeast greenland. earth and planetary science letters 272, 8–18. ó cofaigh, c. & dowdeswell, j.a. 2001: laminated sediments in glacimarine environments: diagnostic criteria for their interpretation. quaternary science reviews 20, 1411–1436. rignot, e. & kanagaratnam, p. 2006: changes in the velocity structure of the greenland ice sheet. science 311, 986–990. roberts, d.h., long, a.j., schnabel, c., freeman, s.p.h.t. & simpson, m.j.r. 2008: the deglacial history of southeast sector of the greenland ice sheet during the last glacial maximum. quaternary science reviews 27, 1505–1516. smith, l.m. & andrews, j.t. 2000: sediment characteristics in iceberg dominated fjords, kangerlussuaq region, east greenland. sedimentary geology 130, 11–25. stein, r., nam, s.i., grobe, h. & hubberten, h. 1996: late quaternary glacial history and short-term ice-rafted debris fluctuations along the east greenland continental margin. in: andrews, j.t. et al. (eds): late quaternary paleoceanography of the north atlantic margins. geological society special publications (london) 111, 135–151. straneo, f., hamilton, g.s., sutherland, d.a., stearns, l.a., davidson, f., hammill, m.o., stenson, g.b. & rosing-asvid, a. 2010: rapid circulation of warm subtropical waters in a major glacial fjord in east greenland. nature geoscience 3, 182–186. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: csa@geus.dk research article colgan et al. 2022: geus bulletin 49. 8315. https://doi.org/10.34194/geusb.v49.8315 1 of 16 sea-level rise in denmark: paleo context, recent projections and policy implications william colgan*1 , hans jørgen henriksen1 , ole bennike2 , sofia ribeiro1 , marie keiding1 , ida karlsson seidenfaden1 , morten graversgaard3 , anne gravsholt busck4 , mikkel fruergaard3 , michael helt knudsen4 , john hopper1 , torben sonnenborg1 , maria rebekka skjerbæk1, anders anker bjørk4 , holger steffen5 , lev tarasov6, r. steven nerem7 , kristian k. kjeldsen1 1geological survey of denmark and greenland (geus), copenhagen, denmark. 2geological survey of denmark and greenland (geus), aarhus, denmark. 3aarhus university, aarhus, denmark. 4university of copenhagen, copenhagen, denmark. 5lantmäteriet, gävle, sweden. 6memorial university of newfoundland, saint john’s, canada. 7university of colorado, boulder, usa abstract we present the most recent intergovernmental panel on climate change sixth assessment report (ar6) sea-level projections for four danish cities (aarhus, copenhagen, esbjerg and hirtshals) under the shared socioeconomic pathway (ssp) family of climate scenarios. these sea-level changes projected over the next century are up to an order of magnitude larger than those observed over the previous century. at these cities, year 2150 sea-level changes of between 29 and 55 cm are projected under the very low emissions scenario (ssp1-1.9), whilst changes of between 99 and 123 cm are projected under the very high emissions scenario (ssp5-8.5). these differences highlight the potentially significant impact of remaining opportunities for climate change mitigation. due to this increase in mean sea level, the mean recurrence time between historically extreme events is expected to decrease. under the very high emissions scenario, the historical 100-year storm flood event will become a 1to 5-year event at most danish harbours by 2100. there is considerable uncertainty associated with these sea-level projections, primarily driven by uncertainty in the future evolution of the antarctic ice sheet and future sterodynamic changes in ocean volume. the ar6 characterises collapse of the west antarctic ice sheet as a low-probability but high-impact event that could cause several metres of sea-level rise around denmark by 2150. in climate adaptation policy, the scientific landscape is shifting fast. there has been a tremendous proliferation of diverse sea-level projections in recent years, with the most relevant planning target for denmark increasing c. 50 cm in the past two decades. translating sea-level rise projections into planning targets requires value judgments about acceptable sea-level risk that depend on local geography, planning timeline and climate pathway. this highlights the need for an overarching national sea-level adaptation plan to ensure municipal plans conform to risk and action standards. *correspondence: wic@geus.dk received: 30 mar 2022 accepted: 10 aug 2022 published: 05 oct 2022 keywords: projection, denmark, coast, sea level, climate scenario abbreviations: ar6: sixth assessment report cmip6: coupled model intercomparison project phase 6 dvr90: 1990 danish vertical reference system ipcc: intergovernmental panel on climate change mis: marine isotope stage psmsl: permanent service for mean sea level rcp: representative concentration pathway ssp: shared socioeconomic pathway geus bulletin is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: marit-solveig seidenkrantz (aarhus university, denmark) reviewed by: three anonymous reviewers funding: see page 14 competing interests: see page 14 additional files: see page 14 introduction future sea-level rise, which threatens coastal populations, ecosystems and infrastructures around the world, is highly dependent on society’s choice of future climate pathway (nauels et al. 2017). denmark is a relatively low-lying nation — approximately 24% of the country lies within 10 m of sea level — with over 7000 km of coastline (fig. 1). presently, approximately 17% of the danish population lives within 6 m of sea level (fryd & jørgensen 2020). this proportion is expected to increase with future population growth focused in coastal cities. the danish coastline is dynamic and has been dramatically shaped by changes in sea level and sediment transport over the past 2000 years. in northern jylland, at skagen, relative sea level has fallen c. 4 m in the past two millennia (hauerbach 1992; clemmensen et al. 2001). in central denmark, near aarhus, as well as offshore in kattegat and øresund, relative sea level has fallen <1 m during the past 2000 years (bennike et al. 2012, 2021; clemmensen et al. 2012; hede et al. 2015). in western jylland, near esbjerg, however, relative sea https://doi.org/10.34194/geusb.v49.8315 https://orcid.org/0000-0001-6334-1660 https://orcid.org/0000-0003-4821-5310 https://orcid.org/0000-0002-5486-9946 https://orcid.org/0000-0003-0672-9161 https://orcid.org/0000-0002-2933-0199 https://orcid.org/0000-0002-7033-1337 https://orcid.org/0000-0001-7636-4335 https://orcid.org/0000-0002-7328-795x https://orcid.org/0000-0002-8575-157x https://orcid.org/0000-0002-7825-695x https://orcid.org/0000-0003-3188-7583 https://orcid.org/0000-0001-9042-8702 https://orcid.org/0000-0002-4919-792x https://orcid.org/0000-0001-6682-6209 https://orcid.org/0000-0003-4064-8111 https://orcid.org/0000-0002-8557-5131 mailto:wic@geus.dk colgan et al. 2022: geus bulletin 49. 8315. https://doi.org/10.34194/geusb.v49.8315 2 of 16 www.geusbul let in.org level has conversely risen c. 1.5 m during the same time (gehrels et al. 2006; szkornik et al. 2008). these contrasting relative sea-level changes, reconstructed over the past two millennia, primarily reflect spatial variations in the glacial isostatic adjustment of earth’s crust associated with deglaciation of the scandinavian ice sheet. on longer timescales, the magnitude of sea-level change is even larger, reflecting absolute increases in ocean volume and elevation associated with the deglaciation of land ice. reconstructed sea-level curves indicate >30 m of local sea-level rise at several danish sites during the mid-holocene, between 12 000 and 6000 years ago (bennike & jensen 2011; bennike et  al. 2012, 2019, 2021). coastal currents and associated sediment transport around denmark play a major role in shaping the present-day 13°e 12°e 12°e 11°e 11°e 10°e 10°e 9°e 9°e8°e 57°n 57°n 56°n 56°n 55°n 55°n < 0 m < 2 m < 10 m 15°e 15°e 55°n 55°n 0 10050 km kattegat ø resund hirtshals esbjerg aarhus copenhagen jylland fyn sjælland m.s.l skagen fig. 1 overview map of denmark with place names referred to in the text. shaded areas designate areas below specified elevations, regardless of whether they have a direct connection to the ocean. this highlights potentially vulnerable areas, should the coastal barrier be breached. the digital elevation map is obtained from the danish agency for data supply and infrastructure. https://doi.org/10.34194/geusb.v49.8315 http://www.geusbulletin.org colgan et al. 2022: geus bulletin 49. 8315. https://doi.org/10.34194/geusb.v49.8315 3 of 16 www.geusbul let in.org shoreline (kabuth et  al. 2014). this underlines a complex history of sea-level change and shoreline position in denmark. global sea level is now once again highly dynamic and increasing, in response to anthropogenic climate change. the rate of 21st century sea-level rise is expected to be faster than that for previous warm holocene intervals, such as the medieval warm period (c. 1000 to 1300 ce) and the holocene thermal maximum (c. 6000 to 4000 bce; gehrels & shennan 2015; kaufman et  al. 2020). paleo studies also infer a consistently high sensitivity of global sea-level rise to atmospheric warming (levermann et  al. 2013). specifically, global sea level was at least 6 m higher than present day during the last interglacial (eemian; marine isotope stage (mis) 5e) when atmospheric temperatures were at least 1°c warmer than pre-industrial temperatures (dutton et al. 2015). the rate of forthcoming anthropocene sea-level rise may be comparable to some of the fastest periods of sea-level rise in the past 12  000  years (weaver et  al. 2003). since 1993, global mean sea level has already risen 8 ± 1 cm (wcrp global sea level budget group 2018). an acceleration in sealevel rise since 1993 is now readily detectable within the global satellite altimetry record, as well as within north sea tide gauge data (nerem et  al. 2018; steffelbauer et  al. 2022). there is tremendous spatial variability in recent vertical land motion and sea-level change across europe (fig. 2). there is a pronounced east–west trend in sea-level change across scandinavia, with sea levels in the gulf of bothnia rising twice as fast as those in the north sea (colgan et al. 2019). the processes influencing recent and future sea-level changes at a specific city are complex. they include larger-scale processes, such as the transfer of land ice into the ocean, the thermal expansion of seawater and shifts in oceanic and atmospheric currents, as well as smaller-scale processes, such as vertical land motion and the redistribution of water masses (nauels et al. 2017). consequently, projected sea-level changes are globally non-uniform, and they vary from city to city (jevrejeva et  al. 2016). here, we examine both the instrumental record of observed sea-level change and the sea-level change projected by the intergovernmental panel on climate change (ipcc) sixth assessment report (ar6), at four danish cities (holgate et  al. 2013; fox-kemper et al. 2021). our aim is to provide policy-relevant insight on local sea-level budgets that resolve all sea-level components, and their associated uncertainties, under the ipcc ar6 shared socioeconomic pathway (ssp) climate scenarios. these scenarios have five fundamental storylines (o’neill et al. 2020; fig. 3). ssp1 characterises a ‘sustainable’ future with very low to low emissions and low challenges associated with climate change adaptation and mitigation. ssp2 characterises a ‘middle of the road’ future with intermediate emissions and challenges. ssp3 characterises a ‘rocky road’ with high emissions and challenges associated with growing regional rivalries. ssp4 characterises a future of ‘inequality’ and high emissions arising from clear regional pathways fig. 2 trends in european ocean and land elevation between january 1993 and december 2020. left: vertical land motion trend observed at fixed global positioning system sites (updated from hammond et al. 2021). right: sea-level elevation trend observed from satellite radar altimetry (updated from nerem et al. 2018). −10° 0° 10° 20° 30° 40° 40° 40° 50° 50° 60° 60° 70° 70° −10 −5 0 5 10 sea level trends (mm/yr) −10° 0° 10° 20° 30° 40° 40° 40° 50° 50° 60° 60° 70° 70° −10 −5 0 5 10 vertical land motion (mm/yr) https://doi.org/10.34194/geusb.v49.8315 http://www.geusbulletin.org colgan et al. 2022: geus bulletin 49. 8315. https://doi.org/10.34194/geusb.v49.8315 4 of 16 www.geusbul let in.org dominated by adaptation challenges. finally, ssp5 characterises a future taking the fast road of ‘fossil-fuel development’ and very high emissions dominated by mitigation challenges. these storylines incorporate broad assumptions about underlying social demographics and offer a wide range of future climates, including multiple alternative pathways to the same climate forcing (fig. 3; meinshausen et al. 2020). the differences in sea-level rise projected under this range of scenarios highlight the significant impact of remaining opportunities for climate change mitigation. methods we employ sea-level projections made available through  the ipcc ar6 sea-level projection tool at https://sealevel.nasa.gov/ipcc-ar6-sea-level-projection-tool (fox-kemper et al. 2021; garner et al. 2021). the ar6 makes sea-level projections for permanent service for mean sea level (psmsl) reporting stations around the world (holgate et al. 2013). within denmark, four psmsl stations are available: aarhus (station 76), esbjerg (station 80), copenhagen (station 82) and hirtshals (station 89; fig. 1). ar6 selected these four stations from the 14 available psmsl stations within denmark to ensure globally homogeneous coverage that is representative of spatial variations in vertical land motion (hansen 2018). at each of these four stations, the sea-level projection is estimated every 10 years between 2020 and 2150 using a local multi-component sea-level budget. the six components of this budget are (1) sterodynamic ocean-volume changes, (2) the antarctic ice sheet, (3) the greenland ice sheet, (4) mountain glaciers, (5) vertical land motion and (6) changes in land-water storage. sea-level projections are made under five illustrative ssp scenarios: ssp1-1.9 (very low emissions), ssp1-2.6 (low emissions), ssp2-4.5 (intermediate emissions), ssp3-7.0 (high emissions) and ssp5-8.5 (very high emissions). the ar6 provides relative sea-level projections as height changes relative to a 2005 baseline. we translate 0 0.5 1 1.5 2 2.5 3 3.5 4 4.5 5 today 2020 g lo ba l-m ea n te m pe ra tu re 2050 1750 2100 ssp1-1.9 ssp1-2.6 ssp2-4.5 ssp3-7.0 ssp5-8.5 ssp5-3.4-os ssp4-6.0 ssp4-3.4 ssp1 sustainability ssp2 middle of the road ssp3 regional rivalry ssp4 inequality ssp5 fossil-fuel development no i t ag i ti m fig. 3 the five ssp storylines with global-mean temperature illustrated relative to pre-industrial levels. the historical temperatures branch into the respective five ssp scenarios over the 21st century. the small black horizontal bars on the 2100 pillars for each ssp indicate illustrative temperature levels. the opaque bands over the 21st century indicate the five ssp scenarios ssp1-1.9, ssp1-2.6, ssp2-4.5, ssp3-7.0 and ssp5-8.5 that are explored in this study. other scenarios (ssp4-6.0, ssp4-3.4 and ssp5-3.4-os) are not explored in this study. modified slightly from meinshausen et al. (2020). https://doi.org/10.34194/geusb.v49.8315 http://www.geusbulletin.org https://sealevel.nasa.gov/ipcc-ar6-sea-level-projection-tool colgan et al. 2022: geus bulletin 49. 8315. https://doi.org/10.34194/geusb.v49.8315 5 of 16 www.geusbul let in.org these sea-level changes to a 1990 baseline by fitting a second-order polynomial curve to the mean annual sealevel elevation observed at each psmsl station between 1900 and 2020. our polynomial curve fitting ignores years with no data. our polynomial approach deviates from the prevailing danish coastal authority approach of applying linear best fit to historical data, which may cause small discrepancies in reported historical sealevel change (ditlevsen et  al. 2019). we estimate the change in sea-level baseline between 1990 and 2005 by differencing the respective annual values of our polynomial fit. this yields 2005 to 1990 baseline corrections to the ar6 relative sea-level changes of +0.2 cm at hirtshals, +1.6 cm at aarhus, +2.0 cm at copenhagen and +2.6 cm at esbjerg (fig. 4). translating these relative projections into absolute sea-level elevation in the standard 1990 danish vertical reference system (dvr90) requires a further correction that reflects the offset between the 1990 baseline sea level and the reported dvr90 elevation for each station (table 1). aside from using a polynomial, rather than linear best fit, this dvr90 conversion conforms to the danish coastal authority approach for linking local dvr90 datum to local sea-level elevation (kystdirektoratet 2018). results the psmsl records at aarhus, esbjerg, copenhagen and hirtshals reveals a spatially complex pattern of sealevel change over the 1900–2020 period, even within the relatively limited size of denmark (fig. 4). hirtshals and esbjerg are clear end members for this sea-level change. at hirtshals, the relative sea level has fallen very slightly since 1900 (–2 ± 2 cm). at esbjerg, by contrast, the relative sea level has increased sharply since 1900 (+16 ± 2 cm). the distance between these two psmsl sites is only c. 255 km. this sharp gradient in recent sea-level change is ultimately associated with the complex pattern of glacial isostatic adjustment following the relatively rapid deglaciation of the scandinavian ice sheet, with denmark becoming ice-free c. 16 000 years ago (houmark-nielsen & kjær 2003). at esbjerg, the earth’s crust is presently subsiding due to the collapse of the crustal forebulge that once ringed the scandinavian ice sheet, whilst at hirtshals, the crust is presently rebounding upwards due to ice unloading (steffen & wu 2011). the post-1900 sea-level changes at copenhagen and aarhus, both +9 ± 2 cm, lie between these end members. the inter-city differences in observed sea-level change are also generally reflected in the projected sea-level table 1 year 1990 elevations of four danish stations (psmsl identification number in parentheses) in the 1990 danish vertical reference system (dvr90). station dvr90 (cm) aarhus (station 76) 6.0 esbjerg (station 80) 20.1 copenhagen (station 82) 10.9 hirtshals (station 89) 1.1 fig. 4 permanent service for mean sea level (psmsl) records of sea-level elevation relative to 1900 at four danish sites (holgate et al. 2013). thick lines are second-order polynomial fits that ignore missing data. these polynomial fits are used to correct the 2005 baseline of ar6 sea-level projections to the 1990 baseline of the danish vertical reference system (dvr90). 15 10 5 0 –5 –10 –15 –20 19 90 se a le ve l( cm ) 1900 1920 1940 1960 1980 2000 2020 year (ce) esbjerg copenhagen aarhus hirtshals https://doi.org/10.34194/geusb.v49.8315 http://www.geusbulletin.org colgan et al. 2022: geus bulletin 49. 8315. https://doi.org/10.34194/geusb.v49.8315 6 of 16 www.geusbul let in.org changes (fig. 5). there can be, however, a non-physical inflection of observed and projected sea-level changes at the year 2020 observation-model transition. this is perhaps most evident at hirtshals and likely results from vertical land motion not being treated consistently between observations and projections. under all climate scenarios, the projected sea-level rise is systematically c. 25 cm higher at esbjerg than at hirtshals. this difference in sealevel rise associated with geography, however, is clearly smaller than the c. 70 cm difference associated with choice of climate scenario. simply put, whilst presently non trivial, with increasing warming, the vertical land motion component of sea-level change becomes small relative to the thermal expansion and cryospheric contributions to sea-level change. for example, at aarhus, the year 2150 sea-level rise ranges from 45 (8–90) cm under the very low emissions scenario (ssp1-1.19) to 115 (69–176) cm under the very high emissions scenario (ssp5-8.5; table 2). here, the 90% confidence intervals indicate that there is a 95% chance that sea-level rise will be less than 90 cm under ssp1-1.9 and less than 176 cm under ssp5-8.5. clearly, however, the sea-level changes observed over the past century are an order of magnitude smaller than the sealevel changes expected over the next century. table 2 presents summary statistics to explicitly address four distinct elements of planning for sea-level change in denmark. first, as discussed earlier, projected sea-level rise varies by location. second, projected sealevel rise varies by year, or planning timeline, of interest. third, projected sea-level rise varies with the choice of climate scenario. finally, and perhaps most importantly, interpreting the appreciable 90% confidence uncertainty is ultimately a risk assessment. ar6 assesses the uncertainty associated with its sea-level projections based on ensemble spread, in both higher-complexity simulations and lower-complexity emulations, as well as structured expert judgment. arguably, the most conservative approach for interpreting an upper-limit metric for sea-level rise planning is pessimistically assuming the very high emissions climate scenario (ssp5-8.5) with 1900 1950 2000 2050 2100 2150 0 50 100 150 200 19 90 se a le ve l( cm ) aarhus 1900 1950 2000 2050 2100 2150 0 50 100 150 200 19 90 se a le ve l( cm ) esbjerg 1900 1950 2000 2050 2100 2150 year (ce) 0 50 100 150 200 19 90 se a le ve l( cm ) copenhagen 1900 1950 2000 2050 2100 2150 year (ce) 0 50 100 150 200 19 90 se a le ve l( cm ) hirtshals ssp1-1.9 ssp2-4.5 ssp5-8.5 observations fig. 5 permanent service for mean sea level (psmsl) observed sea-level change (1900-2020; holgate et al. 2013) and ar6 projected sea-level change (2020–2150; fox-kemper et al. 2021) at four danish cities, expressed relative to a 1990 baseline. shaded areas denote the uncertainty associated with each projection. the ssp1-2.6 and ssp3-7.0 projections are omitted for clarity but appear tabulated in table 2. the c. 2100 ‘bumps’ in projections are a known artifact associated with a decrease in available cmip6 (coupled model intercomparison project phase 6) simulations after 2100, in comparison to before 2100. https://doi.org/10.34194/geusb.v49.8315 http://www.geusbulletin.org colgan et al. 2022: geus bulletin 49. 8315. https://doi.org/10.34194/geusb.v49.8315 7 of 16 www.geusbul let in.org the longest timeline (year 2150) and upper limit (95%) of the confidence interval. this yields upper limit metrics of +186 cm at esbjerg, +179 cm at copenhagen, +176 cm at aarhus and +161 cm at hirtshals. conversely, the ambition of the paris agreement to limit climate change, combined with the need for nearer-term planning targets, can also justify the use of substantially lower planning metrics across denmark. for example, assuming sea-level changes between +22 and +29 cm across all four cities, in the year 2050 under the ssp1-1.9 very low emissions scenario and 95% upper confidence limit. in short, translating the projections in sea-level rise presented here into planning targets requires making value judgements about the acceptable exposures and vulnerabilities to sea-level hazards that depend on geography, timeline and climate choices. discussion projection uncertainty there are two main types of uncertainty associated with the ar6 sea-level projections (fig. 6). the first type of uncertainty is the climate trajectory associated with the five ar6 climate scenarios: ssp1-1.9 (sustainability, very low emissions), ssp1-2.6 (sustainability, low emissions), ssp2-4.5 (middle of the road, intermediate emissions), ssp3-7.0 (regional rivalry, high emissions) and ssp5-8.5 (fossil-fuel development, very high emissions; o’neill et al. 2020). these climate scenarios forecast how society’s greenhouse gas emissions may evolve until 2150. climate scenario uncertainty therefore clearly represents more of a social than physical uncertainty. generally, however, the differences projected between scenarios highlight the significant impact of remaining opportunities for climate change mitigation. simply put, society can still avoid substantial future sea-level rise if we can divert from our current high-emissions climate pathway to a low-emissions climate pathway. the second type of uncertainty is the process-level uncertainty associated with each of the six components in the ar6 sea-level budget: (1) sterodynamic ocean-volume changes, (2) antarctic ice sheet, (3) greenland ice sheet, (4) mountain glaciers, (5) vertical land motion and (6) changes in global land-water storage. both the magnitude and uncertainty of these latter two terrestrial processes are relatively small for all climate scenarios. thus, terrestrial processes are not the main drivers of the sea-level changes projected around denmark by the ar6. next, we describe how other components of the sea-level rise budget influence denmark. the greenland ice sheet contribution and associated uncertainty are an order-of-magnitude smaller than that of mountain glaciers across all five climate scenarios. this is perhaps counter-intuitive, as the greenland ice sheet is a larger contributor to global mean sealevel rise than mountain glaciers. this apparent inversion, however, is an anticipated local consequence of shifting planetary gravity due to the redistribution of water and ice masses (larour et al. 2017). denmark is located within the regional gravitational weakening, and sea-level fall, associated with ice loss from greenland. this makes the greenland ice sheet contribution to sea-level rise disproportionately smaller around denmark, in comparison to the global mean (colgan et al. 2019). denmark is conversely located well beyond the analogous regional gravitational weakening associated with antarctic ice loss and, therefore, receives disproportionately more sea-level rise from the antarctic, in comparison to the global mean. danish sea-level projections therefore incorporate substantially more uncertainty associated with antarctic ice loss, in comparison table 2 the sea-level changes, relative to a 1990 baseline, projected by the ar6 at four danish cities under the five ar6 climate scenarios. climate scenario year esbjerg (cm) aarhus (cm) copenhagen (cm) hirtshals (cm) ssp1-1.9 2050 2100 2150 17 (7‒29) 37 (16‒63) 55 (19‒97) 14 (2‒28) 30 (7‒57) 45 (8‒90) 15 (5‒27) 31 (10‒56) 46 (11‒88) 9 (–3 to +22) 20 (–4 to +47) 29 (–9 to +74) ssp1-2.6 2050 2100 2150 20 (11‒30) 42 (23‒65) 58 (26‒99) 17 (7‒27) 36 (18‒58) 49 (21‒87) 18 (9‒28) 37 (19‒58) 50 (21‒87) 11 (1‒22) 26 (8‒48) 35 (5‒73) ssp2-4.5 2050 2100 2150 22 (13‒32) 55 (37‒78) 85 (51‒131) 19 (10‒29) 49 (32‒72) 77 (44‒121) 20 (11‒29) 50 (33‒72) 78 (46‒122) 13 (4‒23) 38 (21‒61) 61 (28‒106) ssp3-7.0 2050 2100 2150 22 (12‒32) 64 (42‒92) 105 (64‒159) 19 (9‒29) 58 (36‒86) 97 (55‒150) 20 (10‒30) 59 (38‒86) 97 (58‒150) 13 (3‒24) 47 (26‒75) 80 (39‒134) ssp5-8.5 2050 2100 2150 24 (15‒35) 75 (52‒106) 123 (77‒186) 21 (11‒32) 70 (46‒100) 115 (69‒176) 22 (13‒33) 71 (48‒102) 117 (71‒179) 15 (6‒26) 59 (35‒90) 99 (52‒161) the 90% confidence interval is shown in parentheses. within each city and climate scenario, projections are shown for 2050, 2100 and 2150. these relative changes can be translated into absolute dvr90 elevations by adding the city-specific corrections described in the methods. https://doi.org/10.34194/geusb.v49.8315 http://www.geusbulletin.org colgan et al. 2022: geus bulletin 49. 8315. https://doi.org/10.34194/geusb.v49.8315 8 of 16 www.geusbul let in.org esbjerg ssp1-1.9 sea-level contribution (cm) 1 0.8 0.6 0.4 0.2 –40 –20 0 20 40 60 80 100 1 0.8 0.6 0.4 0.2 –40 –20 0 20 40 60 80 100 1 0.8 0.6 0.4 0.2 –40 –20 0 20 40 60 80 100 1 0.8 0.6 0.4 0.2 –40 –20 0 20 40 60 80 100 1 0.8 0.6 0.4 0.2 –40 –20 0 20 40 60 80 100 r el at iv e pr ob ab ili ty (– ) r el at iv e pr ob ab ili ty (– ) r el at iv e pr ob ab ili ty (– ) r el at iv e pr ob ab ili ty (– ) r el at iv e pr ob ab ili ty (– ) ssp1-2.6 sea-level contribution (cm) ssp2-4.5 sea-level contribution (cm) ssp3-7.0 sea-level contribution (cm) ssp5-8.5 sea-level contribution (cm) sterodynamic antarctic is glaciers greenland is vertical land motion land-water storage fig. 6 relative probability distributions of the six contributing terms for the ‘medium confidence’ year 2150 ar6 sea-level budget at esbjerg (fox-kemper et al. 2021). the six components of sea-level change are shown under the five ar6 climate scenarios. the distributions shown here are broadly representative of other danish cities. https://doi.org/10.34194/geusb.v49.8315 http://www.geusbulletin.org colgan et al. 2022: geus bulletin 49. 8315. https://doi.org/10.34194/geusb.v49.8315 9 of 16 www.geusbul let in.org to greenland ice loss. antarctic ice loss, which is a key source of uncertainty in both danish and global projections, is discussed in section 4.2. there is also tremendous uncertainty associated with sterodynamic ocean-volume changes. these are the changes in ocean volume associated with the warming and expansion of ocean water, as well as the dynamic redistribution of ocean mass associated with changing currents. across all climate scenarios, sterodynamic changes are one of the largest components within the ar6 sea-level budget, accounting for 30–40% of the total global sea-level rise, and within every climate scenario, there is substantial uncertainty (year 2150 standard deviation approximately ± 14 cm across the danish sites; fox-kemper et  al. 2021). whilst the relation between water temperature and water density is well-known, and recent thermal expansion is well-observed, modelling the future thermal expansion of the ocean is challenging due to poor scientific understanding of how atmospheric heat is mixed into the near-surface ocean (church et  al. 1991; zanna et  al. 2019; marti et  al. 2022). so how deep will the ocean warm and expand due to spatially and temporally variable processes of downward heat advection and diffusion? the uppermost 700 m of the oceans has absorbed the vast majority of the anthropogenic climate energy associated with global warming since the industrial revolution, but how will the deeper oceans respond to future forcing? in the ar6 projections, the sterodynamic contribution is relatively large across all climate scenarios, which reflects the substantial future thermal expansion committed by past global warming (ehlert & zickfeld 2018). ice-sheet collapse aside from sterodynamic ocean-volume changes, a second critical source of uncertainty in the ar6 projections is the antarctic ice sheet contribution to sea-level rise. the projections described in section 4.1 are regarded as ‘medium confidence’ by the ar6. this is a qualitative term used by the ipcc to convey medium validity based on the type, amount, quality and consistency of data. the ar6 also includes a ‘low confidence’ projection, which, by definition, conveys lower validity from underlying data that incorporates substantially higher near-term contributions from collapsing ice sheets. in comparison to the ‘medium confidence’ projection, which estimates a 5% chance of >230 cm of sea-level rise at copenhagen in 2150, the ‘low confidence’ projection with additional icesheet collapse projects >450 cm at copenhagen in 2150 (fig. 7). whilst the physical mechanisms associated with near-term ice-sheet collapse are qualitatively described, they are difficult to quantify in numerical models (larour et al. 2021). the likelihood of such low probability, but high impact, events is, therefore, inherently difficult to assess beyond expert elicitation (bamber et al. 2019). the primary concern for ice-sheet collapse is the marine ice-sheet instability hypothesis, which fundamentally suggests that the retreat of glaciers grounded fig. 7 ar6 ssp5-8.5 projections with (‘low confidence’) and without (‘medium confidence’) ice-sheet collapse at copenhagen (fox-kemper et al. 2021). the projections shown here are broadly representative of other danish cities. 1900 1950 2000 2050 2100 2150 year (ce) 0 50 100 150 200 250 300 350 400 450 500 19 90 se a le ve l( cm ) copenhagen ssp5-8.5 low ssp5-8.5 medium observations https://doi.org/10.34194/geusb.v49.8315 http://www.geusbulletin.org colgan et al. 2022: geus bulletin 49. 8315. https://doi.org/10.34194/geusb.v49.8315 10 of 16 www.geusbul let in.org below sea level on retrograde bed slopes is irreversible on human time scales (mercer 1978). for the west antarctic ice sheet especially, which is largely grounded below sea level, this means that the retreat of major outlet glaciers can initiate a collapse of multi-metre sea-level equivalent over relatively short time scales (bamber et  al. 2009). there are growing observations of glacier retreat, thinning and acceleration that together suggest that a tipping-point threshold has been crossed, and ice-sheet collapse is already underway in the amundsen sea sector of west antarctica (joughin et al. 2014; rignot et al. 2014; voosen 2021). in a recently structured expert judgment from 23 global sea-level researchers, 5% of respondents expected west antarctica’s sea-level contribution over the 21st century to track a ‘high-end’ (i.e. high emissions) contribution consistent with marine ice-sheet instability (bamber et al. 2019). in the longer term, numerical ice-sheet projections under an ssp5-8.5 very high emissions scenario suggest that both the greenland and antarctic ice sheets are poised to deliver multi-metre sea-level rise for centuries to come (fig. 8; aschwanden et  al. 2019; chambers et al. 2021). more specifically, the ‘high-end’ ensemble members within these projections suggest that complete loss of the greenland ice sheet (c. 7.2 m sea-level equivalent) and west antarctic ice sheet (c. 3.3 m sea-level equivalent) is possible by c. 2500 with currently included model processes. this would raise global mean sea level by >10 m in five centuries, which may be up to half the deglaciation rate of meltwater pulse 1a, which occurred c. 14 500 years ago (weaver et al. 2003; liu et al. 2016). whilst recently observed icesheet mass loss closely tracks the consensus rcp (representative concentration pathway) 8.5 high emissions projection of the recent ipcc ar5 (the imbie team 2020), there is emerging evidence that numerical icesheet simulations systematically underestimate the ice loss observed in the recent past (aschwanden et al. 2021). ice-sheet models can suffer from uncertainty associated with boundary conditions, initial conditions, spatial and temporal resolution, parametric calibration and process inclusion. community assessments derived from an ensemble of ‘best estimate’ models rarely probe these deeper inter-model uncertainties in a systematic fashion (levermann et al. 2014). generally, the scientific debate surrounding ice-sheet collapse has now moved beyond whether it is theoretically possible to whether it is realistically possible under sustained very high emissions. extreme events whilst we present and describe forecasted changes in mean sea level under differing climate scenarios, it is important to recall that local sea level can vary fig. 8 reconstruction of global mean sea-level since the last glacial maximum, spanning the latter part of the pleistocene and the holocene, with the grey dots representing individual observations and black line representing trend after fleming et al. (1998). anthropocene high-emission (rcp8.5/ssp58.5) projections of greenland ice sheet (‘gis’; aschwanden et al. 2019) and west antarctic ice sheet (‘wais’; chambers et al. 2021) collapse are shown, ignoring associated sterodynamic effects, with the ‘high-end’ ensemble member denoted. –20 –15 –10 –5 0 year (1000s ce) –140 –120 –100 –80 –60 –40 –20 0 20 r el at iv e se a le ve l( m ) gis (achwanden et al. 2019) wais (chambers et al. 2021) high-end ensemble pleistocene holocene an tro po ce ne meltwater pulse 1a https://doi.org/10.34194/geusb.v49.8315 http://www.geusbulletin.org colgan et al. 2022: geus bulletin 49. 8315. https://doi.org/10.34194/geusb.v49.8315 11 of 16 www.geusbul let in.org substantially on daily timescales due to tidal and storm effects (wahl et al. 2017; thompson et al. 2021). the 1872 baltic storm, which has been characterised as a once in 4500-year event, delivered extreme sea levels that were up to 280 cm above historical mean sea level in southeast denmark (clemmensen et al. 2014). understanding extreme sea-level events, such as high tide or storm flooding, is critical to understand the inherent risk associated with sea-level rise. the risk of an extreme sealevel event is generally characterised by its recurrence interval. tide gauge records are used to assess the magnitude of historical extreme events, for example characterising the height above mean sea level of the 100-year or 10-year events. as mean sea level around denmark increases, however, the height of extreme sea-level events will also increase, which effectively decreases historical recurrence intervals (su et al. 2021). at aarhus, the 100-year extreme sea-level event has been estimated as +163 cm above mean sea level, based on tide gauge observations during the 1981–2010 period (fig. 9; kystdirektoratet 2018; thejll et al. 2021). under the rcp8.5 high-emissions climate scenario, however, increases in mean sea level result in this same sea-level height recurring as a 5-year event during the 2071–2100 period. su et al. (2021) suggest it will become a 1-year event. in short, sea levels equivalent to the historical 100-year event will occur at least 20 times more frequently during the last decades of the century under the rcp8.5 scenario. as a consequence, the annual duration of high water-level warnings at aarhus will increase from c. 3 h per year in 1981–2010 to c. 1280 h per year (or c. 15% of the year) in 2071–2100 under the rcp8.5 high-emissions climate scenario median estimates (thejll et al. 2021). this general pattern, characterised by decreasing extreme event recurrences and increasing duration of water-level warnings, is similar for other cities around denmark (kystdirektoratet 2018, 2022). these decreases in recurrence intervals are based on stationary statistics and ignore any changes in recurrence due to potential changes in storminess or seasonal cycle. they simply assume that future flooding events will be a composite of historical-type storm surges overlaid on projected sea-level rise. the risk presented by these transient extreme sealevel events is currently managed with a variety of hard and soft solutions around denmark (fig. 10; mangor et al. 2017). hard solutions are engineered constructions of dikes, sluices, sea walls, floodgates and quay–terrain elevation, often with the primary purpose of maintaining the coastline position without any significant changes. soft solutions typically use nature-based approaches, such as establishing new artificial islands, sand dunes or reefs and sand nourishment and/or revegetation along the coastline, with the primary purpose of reducing flood risk. both types of solutions can play a key role in protecting existing and future developments and assets from coastal erosion and inundation (faragó et al. 2018; fryd & jørgensen 2019). the choice between various hard and soft coastal management strategies is ultimately dependent on the municipal management strategy, which can range from the high-burden of ‘holdthe-line’ to low-burden of unmitigated retreat, as well as coastal morphology and the dominant coastal processes during extreme sea-level events. municipal planning we provide a summary of previously published sea-level projections that are relevant for sea-level rise planning at copenhagen in fig. 11. these previous projections, which discount any scenarios of ice-sheet collapse, are collected from studies and assessments published since the 2007 ipcc fourth assessment report (table 3). all these sea-level projections correspond to a ‘high’ or ‘very-high’ emissions climate-change scenario (i.e. the previously published a1b, rcp8.5 scenarios or ssp58.5). the ar6 sea-level projections discussed here are the most recent projections available. whilst previous ipcc assessments reported sea-level rise projections as tabulated global averages, ar6 now provides city-level sea-level projections. after accounting for the use of differing baseline years, these nine studies suggest that the best estimate, for planning purposes, of sea-level rise at copenhagen increased c. 25 cm per decade between 2007 and 2021. there is, however, tremendous variability between studies. for example, the danish climate atlas (thejll et al. 2021) projects 22 cm less sea-level rise in the year 2100 than the ar6 (fox-kemper et al. 2021). fig. 9 storm flood height above mean sea level (cm) for 1-, 5-, 20-, 50 and 100-year recurrence intervals at aarhus. calculated over different 30-year climatology periods until 2100 under the rcp8.5 high emissions climate scenario (kystdirektoratet 2018; thejll et al. 2021). aarhus year (ce) st or m flo od (c m ) 100-year 50-year 20-year 5-year 1-year 20001980 2020 2040 2060 2080 2100 100 120 140 160 180 200 220 https://doi.org/10.34194/geusb.v49.8315 http://www.geusbulletin.org colgan et al. 2022: geus bulletin 49. 8315. https://doi.org/10.34194/geusb.v49.8315 12 of 16 www.geusbul let in.org additionally, these previous studies do not use the same confidence or uncertainty limits, which can make ‘high’ and ‘low’ estimates not directly comparable (table 3). this underlines the need to continually update datasets and projections and make these available for planners, stakeholders and policy makers. there are several reasons for the increasing trend in forecasted sea-level rise. part of this trend reflects an extension of projections from year 2100 to year 2150 in ar6. as sea level rises through time, longer projections yield larger increases in relative sea level. ar6 projects copenhagen-specific sea-level rise to be within c. 10% of the global mean projection. this suggests that adopting global projections for copenhagen-specific planning purposes is generally valid with an additional uncertainty of ± 10% relative to global projections. the shift from more global to more copenhagen-specific projections over the past decade therefore does not explain the worsening sea-level rise forecast at copenhagen. we instead suggest that the increasing trend of sea-level rise projections is primarily due to improved process-level understanding of the influence of climate change on sea-level change over time. it is reasonable to expect this trend to continue. there are diverse sea-level adaptation challenges confronting different coastal municipalities around denmark. the nature of these challenges depends on city layout and topography, as well as inland hydrological conditions (i.e. groundwater level) or coastal hydrodynamic conditions (i.e. tides and waves; mangor et  al. 2017). for example, a city situated within a fjord is confronted with a different flooding hazard than a city situated along an exposed shoreline. consequently, the adaptation challenges confronting danish fig. 10 diversity of management of options for extreme sea-level events. a: coastal wetland in roskilde fjord, sjælland. b: earthen coastal protection dams near rødby, sjælland. c: hardened shoreline at copenhagen, sjælland. d: high-water river lock on ribe å, jylland. c a b d fig. 11 previously published sea-level projections of relevance for copenhagen sea-level rise planning under a high-emissions climate change scenario. all studies expressed relative to 1990 sea level. data labels are associated with specific previous studies in table 3. marker shape indicates global versus copenhagen-specific projections. marker colour indicates year 2100 versus 2150 projections. dashed line denotes the sea-level projection for planning purposes increasing at c. 2.5 cm/yr. the trend shown here is broadly representative of other danish cities. y = 2.5x + c 0 50 100 150 200 250 2005 2010 2015 2020 2025 19 90 se a le ve l( cm ) year 1 2 3 4 5 6 7 8 9 9 global estimate copenhagen estimate year 2100 projection year 2150 projection https://doi.org/10.34194/geusb.v49.8315 http://www.geusbulletin.org colgan et al. 2022: geus bulletin 49. 8315. https://doi.org/10.34194/geusb.v49.8315 13 of 16 www.geusbul let in.org municipalities span across sea-water flooding, river-water flooding, salt-water intrusion and rising groundwater levels. this highlights the need for site-specific solutions and adaptation strategies. in denmark, sealevel rise planning has been decentralised since 2008, with at-risk coastal local municipalities each responsible for their own analyses of flood risks and developing prioritised climate change adaptation action plans. since 2015, eu regulations have also required at-risk coastal municipalities to act to reduce flooding risks (european parliament 2007). madsen et al. (2019) found that whilst danish municipalities are, indeed, aware that rising sea levels will impact their infrastructure, precise municipal responses often remain poorly defined. ergo, there is no national requirement to enforce consistent definitions or updates of municipal sea-level rise adaptation plans. this absence of sea-level rise adaptation plan standards creates the possibility for very different risk management and adaptation actions to occur in nearby municipalities with similar coastal geomorphology and coastal processes. it is evident that local adaptation plans must make critical choices, considering both site-specific hydrology challenges and the general issues of geography, timeline, climate choices and acceptable risk level. arguably, however, there is a pressing need for an overarching common national sea-level adaptation plan to ensure that individual municipal plans conform to some standards in terms of risk analysis using systematically defined vulnerabilities, exposures and hazards. conclusions here, we review sea-level changes at four danish cities during both a past instrumental period and a future projection period. whilst these cities have experienced significant and differing sea-level changes over the instrumental period, the sea-level changes now confronting denmark over the next two centuries are equivalent, or larger, than the sea-level changes experienced in the past two millennia. in a paleo context, the rate of forthcoming anthropocene sea-level rise will likely be comparable to some of the fastest periods of sea-level rise during the holocene, meaning the past c. 12  000 years. translating the local sea-level rise projections presented here into planning targets, however, is challenging. this requires making value judgments about acceptable vulnerabilities and exposures to sea-level hazards that depend on geography, timeline, societal and climatic pathways and risk level. the uncertainty associated with society’s 21st century choice of ssp imparts a tremendous uncertainty on local projections of sea-level rise. the differences in sea-level rise projected under the range of scenarios highlight the potentially significant impact of remaining opportunities for climate-change mitigation. aside from this social uncertainty, across all ssp scenarios, the largest physical uncertainty in projections is associated with the sterodynamic component of sea-level rise. whilst denmark is geographically closer to the greenland ice sheet, uncertainty associated with the antarctic ice sheet’s response to climate change represents a table 3 summary of previously published sea-level projections of relevance for copenhagen sea-level rise planning under a high-emissions climate change scenario. data label study name study year projection baseline projection year projection type scenario low (cm) mean (cm) high (cm) 1 ar4 2007 1990 2100 g a1b +ice sheets 20 41 61 2 swipa2011 2011 1990 2100 g a1b 90 125 160 3 ar5 2013 1996 2100 g rcp8.5 52 74 98 4 jevrejeva2016 2016 2005 2100 c rcp8.5 (+5°c) 55 96 193 5 swipa2017 2017 2005 2100 g rcp8.5 46 74 102 6 klimaatlas 2019 1996 2100 c rcp8.5 6 50 94 7 colgan2019 2019 1925 2100 c rcp8.5 49 77 105 8 srocc 2019 1996 2100 g rcp8.5 61 84 110 9 ar6 2021 2005 2100 g ssp5-8.5 (medium confidence) 63 77 101 9 ar6 2021 2005 2100 c ssp5-8.5 (medium confidence) 48 71 102 9 ar6 2021 2005 2150 g ssp5-8.5 (medium confidence) 98 132 188 9 ar6 2021 2005 2150 c ssp5-8.5 (medium confidence) 51 118 230 data labels correspond to fig. 11 and literature references below the table. g: global; c: copenhagen-specific. all studies are graphically depicted relative to a common 1990 baseline sea level in fig. 4. data labels refer to the following studies: 1: meehl et al. (2007). 2: amap (2011). 3: church et al. (2013). 4: jevrejeva et al. (2016). 5: amap (2017). 6: thejll et al. (2021). 7: colgan et al. (2019). 8: oppenheimer et al. (2019). 9: fox-kemper et al. (2021). https://doi.org/10.34194/geusb.v49.8315 http://www.geusbulletin.org colgan et al. 2022: geus bulletin 49. 8315. https://doi.org/10.34194/geusb.v49.8315 14 of 16 www.geusbul let in.org greater sea-level threat to denmark. the ipcc scenarios analysed here, however, do not assess the likelihood or impact of rapid ice-sheet collapse. over longer, multicentennial, time scales, it is difficult to understate the sea-level threat confronting denmark associated with ice-sheet collapse. the ‘low-probability highimpact’ scenario of ice-sheet collapse and c. 10 m sea-level rise by 2500 cannot be excluded under higher emission scenarios. ice-sheet projections suggest that the only way to avoid certain ice-sheet collapse is to avoid a sustained higher emissions climate pathway. the landscape of sea-level information available to danish policymakers is shifting fast. over the past two decades, the most relevant sea-level rise planning threshold for copenhagen has been increasing at 25 cm per decade. this implies a continual need for updating sea-level projections using the state-of-the-art knowledge and translation of these projections to ensure usability for local stakeholders. there are a variety of hard and soft coastal management strategies already in use in denmark today. the choice of these strategies is ultimately dependent on both municipal management strategy and coastal morphology or processes. whilst sea-level rise planning has been decentralised to local municipalities since 2008, there now appears to be a clear need for an overarching national sea-level adaptation plan to ensure that individual municipal plans for at-risk coastal municipalities conform to both national and eu risk and action standards. acknowledgements we thank the ipcc ar6 ‘ocean, cryosphere, and sea-level change’ chapter for developing and making the sea-level rise projections available, multiple funding agencies for supporting the development of the projections and the nasa sea-level change team for developing and hosting the ipcc ar6 sea-level projection tool available at https://sealevel.nasa. gov/ipcc-ar6-sea-level-projection-tool. we thank aimée slangen (royal netherlands institute for sea research) and robert kopp (rutgers university) for responding to e-mail queries about the ar6 sea-level rise projections. we also thank three anonymous reviewers whose diverse contributions on a wide range of topics greatly improved this interdisciplinary article. finally, we thank marit-solveig seidenkrantz (aarhus university) for serving as scientific editor on this article. additional information funding statement this work is supported by the geocenter denmark project ‘sea-level rise and coastal flooding in denmark: past, future, and policy’, with geocenter partners at the geological survey of denmark and greenland, the department of geoscience at aarhus university and the department of geosciences and natural resource management at the university of copenhagen. competing 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planning conclusions acknowledgements additional information funding statement competing interests author contributions additional files references tables table 1 year 1990 elevations of four danish stations (psmsl identification number in parentheses) in the 1990 danish vertical reference system (dvr90). table 2 the sea-level changes, relative to a 1990 baseline, projected by the ar6 at four danish cities under the five ar6 climate scenarios. table 3 summary of previously published sea-level projections of relevance for copenhagen sea-level rise planning under a high-emissions climate change scenario. figures fig. 1 overview map of denmark with place names referred to in the text. shaded areas designate areas below specified elevations, regardless of whether they have a direct connection to the ocean. this highlights potentially vulnerable areas, should the coastal barrier be breached. the digital elevation map is obtained from the danish agency for data supply and infrastructure. fig. 2 trends in european ocean and land elevation between january 1993 and december 2020. left: vertical land motion trend observed at fixed global positioning system sites (updated from hammond et al. 2021). right: sea-level elevation trend observed from satellite radar altimetry (updated from nerem et al. 2018). fig. 3 the five ssp storylines with global-mean temperature illustrated relative to pre-industrial levels. the historical temperatures branch into the respective five ssp scenarios over the 21st century. the small black horizontal bars on the 2100 pillars for each ssp indicate illustrative temperature levels. the opaque bands over the 21st century indicate the five ssp scenarios ssp1-1.9, ssp1-2.6, ssp2-4.5, ssp3-7.0 and ssp5-8.5 that are explored in this study. other scenarios (ssp4-6.0, ssp4-3.4 and ssp5-3.4-os) are not explored in this study. modified slightly from meinshausen et al. (2020). fig. 4 permanent service for mean sea level (psmsl) records of sea-level elevation relative to 1900 at four danish sites (holgate et al. 2013). thick lines are second-order polynomial fits that ignore missing data. these polynomial fits are used to correct the 2005 baseline of ar6 sea-level projections to the 1990 baseline of the danish vertical reference system (dvr90). fig. 5 permanent service for mean sea level (psmsl) observed sea-level change (1900-2020; holgate et al. 2013) and ar6 projected sea-level change (2020–2150; fox-kemper et al. 2021) at four danish cities, expressed relative to a 1990 baseline. shaded areas denote the uncertainty associated with each projection. the ssp1-2.6 and ssp3-7.0 projections are omitted for clarity but appear tabulated in table 2. the c. 2100 ‘bumps’ in projections are a known artifact associated with a decrease in available cmip6 (coupled model intercomparison project phase 6) simulations after 2100, in comparison to before 2100. fig. 6 relative probability distributions of the six contributing terms for the ‘medium confidence’ year 2150 ar6 sea-level budget at esbjerg (fox-kemper fig. 7 ar6 ssp5-8.5 projections with (‘low confidence’) and without (‘medium confidence’) ice-sheet collapse at copenhagen (fox-kemper et al. 2021). the projections shown here are broadly representative of other danish cities. fig. 8 reconstruction of global mean sea-level since the last glacial maximum, spanning the latter part of the pleistocene and the holocene, with the grey dots representing individual observations and black line representing trend after fleming et al. (1998). anthropocene high-emission (rcp8.5/ssp5-8.5) projections of greenland ice sheet (‘gis’; aschwanden et al. 2019) and west antarctic ice sheet (‘wais’; chambers et al. 2021) collapse are shown, ignoring associated sterodynamic effects, with the ‘high-end’ ensemble member denoted. fig. 9 storm flood height above mean sea level (cm) for 1-, 5-, 20-, 50and 100-year recurrence intervals at aarhus. calculated over different 30-year climatology periods until 2100 under the rcp8.5 high emissions climate scenario (kystdirektoratet 2018; thejll et al. 2021). fig. 10 diversity of management of options for extreme sea-level events. a: coastal wetland in roskilde fjord, sjælland. b: earthen coastal protection dams near rødby, sjælland. c: hardened shoreline at copenhagen, sjælland. d: high-water river lock on ribe å, jylland. fig. 11 previously published sea-level projections of relevance for copenhagen sea-level rise planning under a high-emissions climate change scenario. all studies expressed relative to 1990 sea level. data labels are associated with specific previous studies in table 3. marker shape indicates global versus copenhagen-specific projections. marker colour indicates year 2100 versus 2150 projections. dashed line denotes the sea-level projection for planning purposes increasing at c. 2.5 cm/yr. the trend shown here is broadly representative of other danish cities. research article olivarius et al. 2023: geus bulletin 55. 8309. https://doi.org/10.34194/geusb.v55.8309 1 of 22 mudstone diagenesis and sandstone provenance in an upper jurassic – lower cretaceous evolving half-graben system, wollaston forland, north-east greenland mette olivarius1 , afsoon m. kazerouni2 , rikke weibel1 , thomas f. kokfelt3 , jussi hovikoski4,5 1department for geo-energy and storage, geological survey of denmark and greenland (geus), copenhagen, denmark; 2geological survey of denmark and greenland (geus), copenhagen, denmark; 3department for mapping and mineral resources, geological survey of denmark and greenland (geus), copenhagen, denmark; 4department for geophysics and sedimentary basins, geological survey of denmark and greenland (geus), copenhagen, denmark; 5geological survey of finland (gtk), espoo, finland abstract the influence of rifting on the composition of kimmeridgian to barremian mudstones from northern wollaston forland, north-east greenland is investigated by petrographic and mineralogical analyses of the brorson halvø-1 and rødryg gen-1 cores, and provenance analysis by zircon u-pb age dating of nearby sandstones. mudstone composition varies systematically as a function of the timing of rifting progression and position in the half-graben depositional system. pyrite primarily precipitated in the early rift to rift climax phases. euhedral pyrite overgrowths on framboids formed only during the rift climax phase (lindemans bugt formation). dolomite is the dominant carbonate cement, except for the sediments deposited in the early waning rift phase (palnatokes bjerg formation) where calcite is dominant, and in the late waning rift phase (stratumbjerg formation) where siderite dominates. the highest-temperature reactions with precipitation of illite, quartz, ankerite and barite signify sediment burial depths of >2 km prior to exhumation. uplift-induced fracturing occurred mainly in the early rift to rift acceleration succession (bernbjerg formation). mudstones in the proximal part of the half-graben (rødryggen-1) include more detrital kaolinite than the distal mudstones (brorson halvø-1), which contain more mixed-layer illite-smectite and illite. vermiculite was deposited only in the proximal part of the basin in the rift climax and waning rift successions. chlorite was deposited proximally and distally during the waning rift phase, though supply began earlier in the distal part. fine-grained sediment in the distal part of the half-graben was therefore probably supplied by axial transport from palaeoproterozoic crystalline rocks and mesoto neoproterozoic metamorphic rocks located to the north and north-west. this agrees with the zircon provenance signature from outcropping sand-rich facies, where zircon grains with u-pb ages of 2.0–1.6 ga are dominant, in addition to common 1.6–0.9 ga ages, and fewer 2.8–2.6 ga and 0.47–0.36 ga ages. *correspondence: mol@geus.dk received: 03 jan 2022 revised: 03 oct 2023 accepted: 03 oct 2023 published: 21 dec 2023 keywords: diagenetic processes, mudstone mineralogy, petrography, provenance analysis, rifting abbreviations: bse: backscattered electrons eds: energy dispersive x-ray spectrometry geus: geological survey of denmark and greenland k-s: kolmogorov-smirnov la-icp-ms: laser ablation inductively coupled plasma mass spectroscopy mds: multidimensional scaling se: secondary electrons sem: scanning electron microscopy toc: total organic carbon xrd: x-ray diffraction geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: jon ineson & jørgen bojesenkoefoed (geus, denmark) reviewed by: chris l. kirkland (curtin university, australia), kevin taylor (the university of manchester, uk) funding: see page 20 competing interests: see page 20 additional files: see page 20 1. introduction the archetypal half-graben setting has an ample sediment supply and can be divided into several sedimentation zones related to proximity to bounding faults and rift evolution (surlyk 1978; gawthorpe & leeder 2000). this subdivision contributes to evaluations of sediment distribution in basins. moreover, when knowledge of structural evolution is combined with information on basement rocks in a hinterland area and sediment delivery systems, changes in sediment composition can be linked to source changes over time. possible trends in sediment composition are not well-known in distal marine half-grabens that may be largely isolated from coarse-grained clastic input by coast-parallel submarine rift shoulders. such a setting results in axial sediment transport and mudstone deposition, particularly in basins detached from the coastal area. therefore, the rødryggen-1 and brorson halvø-1 cores from northern wollaston forland in north-east greenland (fig.  1) were used in this study to investigate the proximal versus distal development of a mudstone-dominated half-graben succession deposited during late jurassic to early cretaceous rifting. organic-rich black shales https://doi.org/10.34194/geusb.v55.8309 https://orcid.org/0000-0003-3853-7543 https://orcid.org/0000-0002-8155-8879 https://orcid.org/0000-0001-6311-2593 https://orcid.org/0000-0003-1941-920x https://orcid.org/0000-0001-6330-8713 mailto:mol@geus.dk https://creativecommons.org/licenses/by/4.0/deed.ast olivarius et al. 2023: geus bulletin 55. 8309. https://doi.org/10.34194/geusb.v55.8309 2 of 22 geusbulletin.org have accumulated in many contemporaneous basins, especially in the northern hemisphere (langrock et al. 2003; mutterlose et al. 2003; rogov et al. 2020), so the findings of this study facilitate the prediction of mudstone composition in such basins. the rødryggen-1 and brorson halvø-1 boreholes were drilled in 2009–2010 to depths of 234.5 and 225.7 m, respectively, covering kimmeridgian to barremian sediments (fig. 2) that represent prolonged mudstone deposition in an evolving half-graben setting. the studied half-graben was bounded by north–south-orientated fault crests delimiting this basin from the rest of the rift system and from the palaeo-coast to the west (surlyk 1978). late jurassic – early cretaceous rifting was widespread along the norwegian–greenland seaway, which connected the proto-arctic to the proto-north atlantic (e.g. stoker et al. 2017). we aim to address two main research questions in this study: (1) how do the different phases of half graben development (i.e. early rift, rift acceleration, rift climax, waning rift) affect the composition of the deposited sediment in a mudstone-dominated environment? (2) to what extent does the sediment composition differ between the two cores that represent the proximal versus distal parts of the half-graben system? 2. geological setting the greenlandic craton consists mainly of crystalline basement, which is exposed at the rim of the greenland ice sheet and in tectonic windows in the caledonides (fig. 1; henriksen et al. 2008; kalsbeek et al. 2008a). the sediment source areas relevant for this study are present within the east greenland caledonian fold belt, which originated from the laurentia–baltica continental collision that took place in late cambrian to early devonian time (mckerrow et al. 2000; smith & rasmussen 2008). the orogenesis caused westwards thrusting of crystalline complexes of archaean to palaeoproterozoic age, metasediments of the krummedal and smallefjord sequences of mesoproterozoic to early neoproterozoic age, metasediments of the eleonore fig. 1 geological map of north-east greenland based on stemmerik et al. (1997), henriksen et al. (2008) and kalsbeek et al. (2008a). detailed map of wollaston forland based on the digital greenland geological map at a scale of 1:500 000 and the printed map series at a scale of 1:100 000. the locations of the rødryggen-1 and brorson halvø-1 boreholes and the samples for zircon u-pb age dating are shown. thrust fault/shear zoneneoproterozoic and caledonian granites neoproterozoic (mainly eleonore bay supergroup) meso–neoproterozoic metamorphic rocks (including krummedal and smalle�ord sequences) archaean–palaeoproterozoic crystalline complex tectonic windows with crystalline rocks devonian cambrian–silurian triassic permian carboniferous jameson land g re e nlan d ice sh e e t kong oscar fjord 100 km 25°w70°n 74°n liverpool land score s b y s und a–g localities for zircon age data from literature (see fig. 9) wollaston forland clavering ø hold with hope traill ø geographical society ø 72°n palaeogene volcanic rocks cretaceous jurassic quaternary palaeogene sills and dykes borehole provenance sample a b a b d e c f d c c b a shannon ø hochstetter forland kuhn ø kejser franz joseph fjord a a payer land g 10 km 20°w 20°w rødryggen-1 fladebugt young sund wollaston forland daneborg brorson halvø-1 123 4 https://doi.org/10.34194/geusb.v55.8309 http://www.geusbulletin.org/ olivarius et al. 2023: geus bulletin 55. 8309. https://doi.org/10.34194/geusb.v55.8309 3 of 22 geusbulletin.org bay supergroup of neoproterozoic age and sediments of neoproterozoic to silurian age (kalsbeek et al. 2000, 2008b; watt et al. 2000; thrane 2002; higgins et al. 2004). caledonian metamorphism, migmatization and granite intrusion took place in ordovician to silurian times, and the crustal thickening resulted in continental sedimentation during the devonian, followed by post-caledonian terrestrial and marine sedimentation in the carboniferous to palaeogene (stemmerik et  al. 1992; kalsbeek et al. 2001; gilotti et al. 2008; larsen et al. 2008). late triassic – middle jurassic thermal subsidence in the sedimentary basins was followed by middle jurassic rifting that was succeeded by transgression in the late jurassic and then renewed rifting, which culminated in the latest jurassic – earliest cretaceous (surlyk 2003). the wollaston forland peninsula provides one of the most complete stratigraphic records of the jurassic and early cretaceous in north-east greenland (fig. 1). the cored section covers the kimmeridgian to lower barremian interval, which is divided into four formations: (1) the bernbjerg formation, (2) the lindemans bugt formation (storsletten member), (3) the palnatokes bjerg formation (albrechts bugt and rødryggen members) and (4) the stratumbjerg formation (fig. 2). the bernbjerg formation spans the late oxfordian to the early volgian and forms an up to 500–600 m thick black mudstone – shale succession that accumulated in a tectonically-affected shelf setting (e.g. surlyk et al. 2021). rifting intensified during the volgian, fragmenting the basin into a series of narrow, 10–30 km wide, westward tilted, fjord-like half-grabens. major conglomeratic submarine fan-delta systems (lindemans bugt formation, rigi member) developed in the most proximal fault block reaching a maximum thickness of 2 km (surlyk 1978; henstra et al. 2016). the coeval palaeoenvironmental development in more distal fault blocks has remained poorly understood due to lack of outcrops but is well-recorded in the studied cores. the new data indicate deep basinal sedimentation (distal part of lindemans bugt formation) and detachment of the permpas–hühnerbjerg blocks (in which the studied boreholes are located) from the coastal deltaic systems (hovikoski et al. 2023a (this volume), b). the rift climax lasted until the valanginian and was followed by waning rift activity and transgression in the western part of the study area (surlyk 1978, 1984, 1990, 2003). an up to 600 m thick succession of gravity-flow deposits with conglomerates and sandstones (palnatokes bjerg formation, young sund member) accumulated in the proximal fault block, whereas fossiliferous mudstones (albrechts bugt and rødryggen members) were deposited in basinal areas and on submarine block crests (surlyk 1978, 1984, 2003; surlyk & korstgård 2013; hovikoski et al. 2018). the rødryggen-1 and brorson halvø-1 cores penetrate both mudstone members. new biostratigraphic data (nøhr-hansen et  al. 2020; alsen et  al. 2023, this volume) suggest a valanginian to hauterivian age for these deposits. towards the east, fault activity continued until the barremian and led to the deposition of the coarse-grained falskebugt member (piasecki et al. 2020). during the late hauterivian, deposits of the palnatokes bjerg formation were drowned and succeeded by sub-storm wave-base bioturbated mudstones of the stratumbjerg formation (bjerager et al. 2020). a several metres thick upper hauterivian to lower barremian interval of the lowermost stratumbjerg formation is recorded at the top of the brorson halvø-1 core (alsen et al. 2023, this volume) and younger parts of the formation of barremian to albian age are present in outcrops near the drill site (piasecki et al. 2020). 3. methodology 3.1 petrography the petrographic and mineralogical characteristics of cemented and laminated mudstones in the fig. 2 stratigraphic scheme with vertical lines showing the rødryggen-1 (rø-1) and brorson halvø-1 (bh-1) cored successions. al: albrechts bugt mb. b: bernbjerg fm. ba: bastians dal fm. j: jakobsstigen fm. l: laugeites ravine mb. li: lindemans bugt formation. li (s): lindemans bugt fm (storsletten mb). mu: muslingebjerg fm. n: niesen mb. pa: palnatokes bjerg fm (young sund mb). pay: payer dal fm. pe: pelion fm. r: rødryggen mb. ri: rigi mb. str: stratumbjerg fm. ug: ugpik ravine mb. modified from bojesen-koefoed et al. 2023a, this volume). s n alluvial/delta plain, paralic, sand-dominated coal shallow marine sandstones shelf transition – sandstones, mudstones, heteroliths o�shore/basinal mudstones deep marine sandstones deep marine conglomerates calcareous sandy marine mudstones red marine mudstones hiatus/condensed chronostratigraphywollaston forland – kuhn ø r ? al str bh-1 rø-1pa no data li li (s)ri n b j pe ug pay mu ba onlaps crystalline basementonlaps upper permian w e barremian hauterivian valanginian ryazanian volgian berriasian tithonian kimmeridgian oxfordian callovian bathonian bajocian aalenian c re ta ce ou s ju ra ss ic lo w er u pp er m id dl e u m u l l u m l u m l u m l u l u m l l ? https://doi.org/10.34194/geusb.v55.8309 http://www.geusbulletin.org/ olivarius et al. 2023: geus bulletin 55. 8309. https://doi.org/10.34194/geusb.v55.8309 4 of 22 geusbulletin.org rødryggen-1 and brorson halvø-1 cores were studied by transmitted and reflected light microscopy as well as by scanning electron microscopy (sem). samples of fracture fillings and sandy intervals were also examined. polished thin sections were prepared from selected intervals in the cores representing diagenetic features characteristic of the different types of lithologies. the thin sections were impregnated with blue epoxy to ease identification of open pore space. half of each thin section was etched and stained with sodium cobaltinitrite to facilitate k-feldspar identification. the sem analyses were performed at the geological survey of denmark and greenland (geus) using a philips xl40 sem equipped with a thermonoran energy dispersive x-ray spectrometry (eds) detector that was used to analyse the chemistry. selected samples were studied with a backscattered electron (bse) detector on carbon-coated thin sections and with a secondary electron (se) detector on gold-coated rock chips. the mineralogy of each type of lithology present in the rødryggen-1 and brorson halvø-1 cores was analysed by x-ray diffraction (xrd) at the university of copenhagen, denmark, and geus. the edge of the core was removed to avoid contamination. the bulk mineralogy was measured on a bruker advance d8 diffractometer with a lynxeye detector using samples crushed to <63 μm applying the bragg–brentano method. semi-quantification of the bulk mineralogy was obtained by the rietveld method (rietveld 1969; mccusker et  al. 1999). the clay fraction analysis was carried out on a philips 1050 goniometer with fixed divergence, anti-scatter slits and co-kα radiation (pulse high selection and fe-filter). chemical pre-treatment with naocl at ph 9.0 was used to remove organic matter. the samples were dispersed ultrasonically in distilled water to acquire the clay fraction (<2 µm). the >30 µm fraction was removed by density separation and the intermediate fraction by centrifugation in a centrifugal particle size analyser (slater & cohen 1962). the suspensions were flocculated in 1 m nacl, and excess salt was removed by centrifugation and washing with water and ethanol. three orientated specimens were made for each sample by the pipette method, comprising mg-saturated air-dry, mg-saturated with glycerol, and k-saturated air-dry heated at 300°c for 1 h. an x-ray diffractogram was produced for each of the saturated specimens on which the discrete minerals were identified from peak positions (hillier 2000) and semi-quantified by application of correction factors. 3.2 zircon u-pb geochronology the rødryggen-1 and brorson halvø-1 cores did not contain sufficiently coarse material to apply detrital zircon u-pb age dating methods. instead, outcrop samples collected from three locations east of the rødryggen-1 drill site and a location situated north-east of the brorson halvø-1 drill site were used for the provenance analysis (fig.  1). the samples collected near the rødryggen-1 drill site consist of silty sandstone to sandy siltstone belonging to the bernbjerg formation, corresponding to the lower part of the cored succession. the sandstone sampled near the brorson halvø-1 drill site belongs to the albian part of the stratum bjerg formation, so it is younger than the part of the formation encountered in the brorson halvø-1 core. the detrital zircon u-pb age analyses were performed by laser ablation inductively coupled plasma mass spectroscopy (la-icp-ms) at geus. samples were crushed and sieved to retrieve the grain-size fraction <500  µm. a water-shaking wilfley table was used to obtain heavy mineral concentrates. zircon grains were hand-picked in a random way to ensure that a range of grain sizes, shapes and colours were included. the polished epoxy mount with the zircon grains was cleaned in an ultrasonic bath with propanol and loaded into the sample cell of the laser ablation system for radiometric age dating. the data were acquired with a single spot analysis on individual zircon grains. a  beam diameter of 30 µm and a crater depth of c. 15–20 µm were used. the  amount of ablated material was c. 200–300 ng for the ablation time of 30 sec. the ablated material was analysed on an element2 (thermo finnigan) single-collector, double focusing, magnetic sector-field, inductively coupled plasma mass spectrometer with a fast-field regulator for increased scanning speed. the total acquisition time was 60 sec for each analysis, of which the first 30 sec were used to measure the gas blank. the instrument was tuned to give large, stable signals for the 206pb and 238u peaks, low background count rates (typically around 150 counts per second for 207pb) and low oxide production rates (238u16o/238u generally below 2.5%). 202hg, 204(pb+hg), 206pb, 207pb, 208pb, 232th and 238u intensities were determined through peak jumping using electrostatic scanning in low resolution mode and with the magnet resting at 202hg. mass 202hg was measured to monitor the 204hg interference on 204pb where the 202hg/204hg ≡ 4.36, which can be used to correct significant common pb contributions using the model pb composition of stacey & kramers (1975). standard-sample bracketing using the gj-1 zircon (jackson et  al. 2004) was used to correct the elemental fractionation induced by the laser ablation and the instrumental mass bias on measured isotopic ratios. long-term external reproducibility was monitored by repeated analyses of the plešovice zircon standard (sláma et  al. 2008). the reported ages are based on https://doi.org/10.34194/geusb.v55.8309 http://www.geusbulletin.org/ olivarius et al. 2023: geus bulletin 55. 8309. https://doi.org/10.34194/geusb.v55.8309 5 of 22 geusbulletin.org 207pb/206pb derived ages for the >0.8 ga (billion years) analyses and 206pb/238u ages for the <0.8 ga analyses, since the latter is more precise for the younger age range, and a natural gap between age populations exists. the propagation of the analytical errors follows the principles of sambridge & lambert (1997). age measurements were discarded if they lacked a stable 207pb/206pb plateau or for u/pb or pb/pb error >10%. a correction for common pb was applied on a small fraction (≤7%) of the concordant analyses from each sample. the data are plotted using kernel density estimation (vermeesch 2012) employing analyses with <10% discordance. the analytical data are reported in supplementary file s1. 4. results 4.1 lithology the studied succession was deposited during a protracted rifting episode that is differentiated into four discrete rift phases: (1) early rifting during the kimmeridgian part of the bernbjerg formation, (2) rift acceleration during the volgian part of the bernbjerg formation, (3) rift climax during deposition of the lindemans bugt formation and (4) waning rifting during deposition of the palnatokes bjerg formation (figs 3 and 4; surlyk 1978, 2003). these major rift phases resulted in significant shifts in depositional conditions that are reflected in the lithological characteristics observed in the different stratigraphic units in the rødryggen-1 and brorson halvø-1 cores. in the cored sections, the bernbjerg formation comprises dark grey mudstones with interlaminated coarse siltstones to very fine sandstones in some intervals (fig. 5a). the lindemans bugt formation consists of dark grey clayey mudstones (fig. 5c) with a larger content of fossils and pyrite than in the bernbjerg formation. the albrechts bugt member of the palnatokes bjerg formation comprises light grey mudstones (fig. 5d) that differ from the lindemans bugt formation in being more calcareous, sandy and bioturbated. the rødryggen member of the palnatokes bjerg formation consists of red hematitic mudstones with intercalated sandy mudstones (fig.  5e). the stratumbjerg formation contains bioturbated grey mudstones (fig.  5f). the sedimentological characteristics are described in more detail by hovikoski et al. (2023a, this volume). carbonate-cemented intervals occur in all the stratigraphic units and are characterised by lower gamma-ray values (figs 3, 4) and lighter colours (fig. 5b). the lighter colours are also evident in the microscopic appearance of the cemented mudstones due to the lower clay mineral content as compared to the uncemented mudstones (figs 6a, 6b). pyrite is evident in many core intervals as well as various macrofossils, deformation structures, faults and fractures (figs 3 and 4). 4.2 detrital components the mineralogy of the rødryggen-1 and brorson halvø-1 cores based on xrd analyses is presented in figs 3 and 4 for the 27 analysed mudstones and two samples of cemented fractures. quartz is the most abundant mineral amounting to 12–48 wt% with an average of 29 wt%, which is present as siltand sand-sized grains in the mudstones (fig. 6c). k-feldspar and plagioclase/albite occurs in amounts of up to 11 and 18 wt%, respectively, and each of them are present as 6 wt% on average. petrographic analysis of the k-feldspar grains reveals that they are generally well-preserved, whereas the plagioclase and albite grains are often partially dissolved (fig.  6d). apatite is found as detrital clasts (fig. 6e) in amounts of 4 wt% on average with the highest contents up to 16 wt% occurring in a condensed interval in the lower part of lindemans bugt formation in the rødryggen-1 core. calcite is present as detrital clasts, including bioclasts in the palnatokes bjerg formation, in both the albrechts bugt and rødryggen members. the bernbjerg and lindemans bugt formations are rich in organic matter including some coal fragments. detrital heavy minerals are found in accessory amounts and comprise primarily ilmenite, leucoxene, rutile, magnetite, zircon and garnet in the form of almandine. muscovite could not be differentiated from illite by xrd but was observed in thin section. the mica minerals primarily consist of muscovite with subordinate biotite. the micas are generally aligned parallel to the lamination and are often cleaved into thin sheets and bend around less ductile grains (fig. 6c). most of the clay minerals are detrital as testified by the absence of growth structures and by their tangential orientation around the other detrital minerals (fig.  6f). kaolinite is present in the matrix of all the mudstones and occurs in higher amounts in the rødryggen-1 core than in the brorson halvø-1 core (figs 3, 4). kaolinite is found in amounts up to 17 wt% with an average of 10 wt%. mixed-layer illite-smectite is found in amounts of 7 wt% on average with the highest contents occurring in the brorson halvø-1 core and especially in the sample from the stratumbjerg formation where it constitutes 28 wt%. illite occurs with an average of 13 wt% and is more abundant overall in the brorson halvø-1 core than in the rødryggen-1 core. the smallest illite contents within each well occur in the palnatokes bjerg formation and the highest contents of up to 25 wt% are present in the bernbjerg formation. some of the illite is authigenic as evident by its morphology, but the proportion of detrital to authigenic illite cannot be quantified. chlorite is found in all samples from the brorson halvø-1 core (fig. 6f), except the two deepest samples from the bernbjerg formation, and its content increases https://doi.org/10.34194/geusb.v55.8309 http://www.geusbulletin.org/ olivarius et al. 2023: geus bulletin 55. 8309. https://doi.org/10.34194/geusb.v55.8309 6 of 22 geusbulletin.org fig. 3 mineralogy from xrd of the rødryggen-1 core plotted with the sedimentological log. chronostratigraphy from alsen et al. (2023, this volume). see fig. 4 for legend. chronostrat.: chronostratigraphy. lithostrat.: lithstratigraphy. gr: gamma ray. api: american petroleum units. b er nb je rg f m li nd em an s bu gt f m w ol la st on f or la nd g ro up h al l b re dn in g g ro up pa ln at ok es b je rg f m a lb re ch ts b ug t m b clay silt vf f m c vc sand gr (api) rødryggen-1chronostrat. lithostrat. tectonic phase ki m m er id gi an lo w er v ol gi an m id dl e v ol gi an u pp er ry az an ia n lo w er r ya za ni an u pp er v ol gi an u pp er va la ng ia n lo w er v al an gi an py py py py an an an an an an py py py py py py py structures and fossils2200 pyri ft c lim ax w an in g ri ft ri ft a cc el er at io n ea rl y ri ft 110 120 130 140 150 160 170 180 190 200 210 220 0 10 20 30 40 50 60 70 80 90 100 230 234.5 m mineralogy (wt%) 0 20 40 60 80 100 2.70 3.30 13.31 22.27 23.88 26.60 57.43 66.45 74.60 76.67 89.62 89.66 154.54 173.51 173.63 194.29 194.41 205.73 221.00 222.67 227.32 fracture �ll m quartz k-feldspar plagioclase apatite calcite ankerite-dolomite pyrite kaolinite mixed-layer clays illite chlorite vermiculite https://doi.org/10.34194/geusb.v55.8309 http://www.geusbulletin.org/ olivarius et al. 2023: geus bulletin 55. 8309. https://doi.org/10.34194/geusb.v55.8309 7 of 22 geusbulletin.org fig. 4 mineralogy from xrd of the brorson halvø-1 core shown alongside the sedimentological log. chronostratigraphy from alsen et al. (2023, this volume). chronostrat.: chronostratigraphy. lithostrat.: lithstratigraphy. gr: gamma ray. api: american petroleum units. py py py py py py py py py py py py py py py mfr mfr ? ? ? py py py py py py py py py py py py py py py py an an an an py py py dominated py py py py py py py py py py brorson halvø-1 225.7 110 120 130 140 150 160 170 180 190 200 210 220 0 10 20 30 40 50 60 70 80 90 100 an l. ba rr em ia n lo w er v ol gi an ki m m er id gi an clay silt vf f m c vc sand u. h. up pe r v al an gi ni an – m id de l h au te ri vi an u. r. m.vol. ? u. v. mchronostrat. lithostrat. tectonic phase structures and fossils2200 b er nb je rg f m h al l b re dn in g g ro up li nd em an s b ug t f mw ol la st on f or la nd g ro up pa ln at ok es b je rg f m a lb re ch ts b ug t m b r ød ry gg en m b st ra tu m bj er g fm b ro rs on h al vø g r. ri ft cl im ax w an in g ri ft ri ft a cc el er at io n ea rl y ri ft m mineralogy (wt%) 0 20 40 60 80 100 3.18 9.70 23.20 35.20 38.15 52.71 117.70 212.70 mfr concretion open fractures clay clast bivalve wave ripple cross-strati�cation coal belemnite contorted lamination shell fragments plant fragments ripple cross-strati�cation faults fractures very �ne sandstone/coarse siltstone interlaminated sandstone and mudstone calcite-cemented sandy mudstone ankerite and dolomite-cemented mudstone planar laminated mudstone mottled lamination pyrite ankerite ammonite onychites slump loading synaeresis crack scour-and-�ll (gutter cast?) mud �occule ripple py an fracture �ll quartz k-feldspar plagioclase apatite calcite ankerite-dolomite hematite pyrite kaolinite mixed-layer clays illite chlorite vermiculite gr (api) https://doi.org/10.34194/geusb.v55.8309 http://www.geusbulletin.org/ olivarius et al. 2023: geus bulletin 55. 8309. https://doi.org/10.34194/geusb.v55.8309 8 of 22 geusbulletin.org upwards with the highest amount of 18 wt% occurring in the stratumbjerg formation (figs 3, 4). in the rødryggen-1 core, chlorite is only present in the palnatokes bjerg formation where it occurs in amounts of 2–4 wt%. vermiculite is not present in the mudstones from the brorson halvø-1 core except for a small content (1 wt%) in the stratumbjerg formation. in the rødryggen-1 core, vermiculite is present in all samples from the lindemans bugt and palnatokes bjerg formations in contents of 4–14 wt%, whereas it is absent in the bernbjerg formation. 4.3 authigenic minerals pyrite is on average 5 wt% and it was the first mineral that precipitated in the sediments. pyrite is present fig. 5 core photos of the mudstone texture of cemented versus uncemented mudstones. bh-1: brorson halvø-1 core. rø-1: rødryggen-1 core. a: laminated sandy mudstone. b: cemented mudstone. c: laminated pyritic mudstone. d: bioturbated mudstone. e: bioclastic hematitic cemented mudstone. f: bioturbated mudstone. a bernbjerg fm, rø-1, ~166 m 1 cm b lindemans bugt fm, storsletten mb, rø-1, ~76.6 m 1 cm f stratumbjerg fm, bh-1, ~2 m 1 cm c lindemans bugt fm, storsletten mb, rø-1, ~64 m 1 cm e palnatokes bjerg fm, rødryggen mb, bh-1, ~20 m 1 cm d palnatokes bjerg fm, albrechts bugt mb, rø-1, ~2.7 m 1 cm https://doi.org/10.34194/geusb.v55.8309 http://www.geusbulletin.org/ olivarius et al. 2023: geus bulletin 55. 8309. https://doi.org/10.34194/geusb.v55.8309 9 of 22 geusbulletin.org in all formations except some parts of the stratumbjerg and palnatokes bjerg formations in the brorson halvø-1 core (figs 3, 4). it is often found in association with organic matter (fig.  7a). the highest amounts (4–14 wt%) are found in the lindemans bugt formation where the pyrite framboids are often overgrown by euhedral pyrite (fig. 7b). only framboidal pyrite is present in the remaining formations where it constitutes 4–7 wt% in the bernbjerg formation and 0–3 wt% in the palnatokes bjerg formation. although the kaolinite crystals have euhedral shapes, they are likely to have been transported before fig. 6 texture and detrital phases. a: typical mudstone texture. b: cemented mudstone texture. c: quartz occurs as siltand sand-sized grains in the mudstones. d: partially dissolved albite grain. e: apatite occurs as detrital clasts. f: most clay minerals are detrital such as chlorite. a bernbjerg fm, tl image, rø-1, 201.81 m b bernbjerg fm, tl image, rø-1, 187.50 m c bernbjerg fm, bse image, bh-1, 212.70 m d palnatokes bjerg fm, bse image, rø-1, 13.31 m e lindemans bugt fm, bse image, rø-1, 89.96 m f stratumbjerg fm, se image, bh-1, 3.18 m 100 μm 100 μm 20 μm50 μm 10 μm 5 μm m q q c f a k d a ap ch k bh-1: brorson halvø-1 core rø-1: rødryggen-1 core tl: transmitted light cn: crossed nicols bse: backscattered electron se: secondary electron a: ankerite ap: apatite b: barite c: calcite ch: chlorite d: dolomite f: feldspar k: kaolinite m: mica ml: mixed-layer clay o: opal om: organic matter pe: pyrite euhedra pf: pyrite framboid q: quartz https://doi.org/10.34194/geusb.v55.8309 http://www.geusbulletin.org/ olivarius et al. 2023: geus bulletin 55. 8309. https://doi.org/10.34194/geusb.v55.8309 10 of 22 geusbulletin.org deposition since they do not occur in well-defined booklets (fig. 7c). however, some detrital grains have been kaolinised, including some of the muscovite and feldspar grains, which must have happened within the sediment since they would have disintegrated during transport. it is often difficult to distinguish unambiguously between the detrital and authigenic clay minerals, but the morphology of the mixed-layer illite-smectite and illite indicate that part of it is authigenic. fibrous and hairy illite was observed using sem. thin quartz overgrowths have mainly precipitated in the coarser-grained intervals (fig. 7c) and overgrow pyrite and kaolinite. barite has precipitated in some of the mudstones where it formed as the last authigenic phase. a bernbjerg fm, bse image, bh-1, 52.71 m b lindemans bugt fm, bse image, rø-1, 76.47 m c bernbjerg fm, se image, rø-1, 227.32 m d lindemans bugt fm, se image, rø-1, 90.01 m e bernbjerg fm, bse image, rø-1, 187.50 m f palnatokes bjerg fm, bse image, rø-1, 23.88 m 20 μm 20 μm 5 μm 5 μm5 μm 30 μm om pf pf pe a d q k a b mi pf a d q c fig. 7 authigenic minerals. a: framboidal pyrite has often formed in connection with organic matter. b: euhedral pyrite has only formed in lindemans bugt formation where it has overgrown pyrite framboids. c: quartz has overgrown detrital kaolinite. d: barite has overgrown ankerite rhombs. e: dolomite is the dominant carbonate cement in bernbjerg and lindemans bugt formations. f: calcite is the dominant carbonate cement in palnatokes bjerg formation. see fig. 6 for abbreviations. https://doi.org/10.34194/geusb.v55.8309 http://www.geusbulletin.org/ olivarius et al. 2023: geus bulletin 55. 8309. https://doi.org/10.34194/geusb.v55.8309 11 of 22 geusbulletin.org it occurs as string-like precipitations and as euhedral crystals overgrowing ankerite (fig. 7d). barite was identified by eds but has formed in such small amounts that it could not be estimated by xrd analysis. pervasive carbonate cementation of the mudstones is found in some intervals of all the studied formations, and these intervals are characterised by low gamma-ray values in the natural gamma log (figs 3, 4). high concentrations of bioclasts or micrite are present in the carbonate-cemented intervals, except for those with siderite. dolomite is the dominant carbonate mineral in the bernbjerg and lindemans bugt formations, whereas calcite is dominant in the palnatokes bjerg formation and siderite is dominant in the stratumbjerg formation. dolomite and ankerite could not be clearly discriminated by xrd so they have been grouped as ankerite– dolomite in the xrd results (figs 3, 4). however, both minerals are present since they were identified by eds. they are found in all formations with the highest abundance in the bernbjerg formation (maximum 34 wt%) followed by lindemans bugt formation (maximum 8 wt%). poikilotopic dolomite occurs in highest abundance, whereas ankerite occurs mostly as smaller rhombs (fig.  7e). ankerite is often precipitated between exfoliated mica flakes. an outwards increase in fe content is observed in both the dolomite and ankerite crystals. calcite occurs in amounts of 2–4 wt% in the samples from the bernbjerg and lindemans bugt formations, reaching 6–54 wt% in the palnatokes bjerg formation comprising both detrital and authigenic calcite that is poikilotopic and micritic (fig.  7f). a calcite content of 1 wt% is found in the sample from the stratumbj erg formation. this sample does not contain siderite, but it is present in other samples from this stratigraphic unit where it has precipitated as an early phase. 4.4 bioclasts the bernbjerg and lindemans bugt formations contain numerous calpionellids that are most abundant in the cemented mudstones. the calpionellids and other shell material in the bernbjerg and lindemans bugt formations are recrystallised, but moulds of molluscs and ammonites have been found (alsen et al. 2023). calpionellids are calcareous microfossils of uncertain affinity. they have oblong shells which have been filled with either dolomite or micritic calcite and often also ankerite and pyrite (fig. 8a). pyrite precipitated mostly along shell rims though sometimes filling most of the internal cavity. ankerite formed small euhedral crystals, most of which precipitated on the exterior of the bioclasts. poikilotopic dolomite crystals precipitated in most of the remaining cavity, but some porosity is often preserved (fig. 8b). fossils in the palnatokes bjerg formation include ostracods, brachiopods, foraminifera and inoceramid bivalves (alsen et  al. 2023). the formation contains abundant calcispheres, especially in the cemented intervals. they have not been recrystallised as seen by the characteristic test (fig. 8c) and by the extinction pattern following the growth structure in other fossils. these calcispheres are probably calcareous dinoflagellate cysts and have a spherical test. calcite has precipitated in the interior, and small ankerite crystals have sometimes formed within the calcite (fig. 8d). 4.5 fractures small fractures are present in most of the core and are most evident in the cemented intervals where they cut through the carbonate-cemented fabrics. opal has often formed along the rims of the fractures where it radiates in multiple layers and forming spherical layers around a protruding matrix (fig. 8e). the remaining parts of the fractures are filled with dolomite in which the fe content decreases towards the middle of the fractures (fig. 8f). the dolomite has occasionally replaced some of the opal along its outer rim (fig. 8e). xrd analyses have been made of fracture fills comprising one sample selected from each well in the bernbjerg formation where fractures are most abundant, and show that dolomite is the dominant fracture-filling cement (figs 3, 4). fracturing has happened several times as seen by the cross-cutting relationships of the fracture generations, where each of them became cemented prior to the next generation of successively wider fractures. the largest encountered fractures are up to a few centimetres wide. the last generation of fractures were not filled by any minerals and thus increased the porosity and permeability. 4.6 zircon u-pb ages the detrital zircon u-pb ages of the four outcrop samples from wollaston forland (fig. 1) cover a broad mesoarchaean to palaeozoic age span (fig. 9). the discordant ages (comprising 21–28%) are not plotted but included in supplementary file s1. the three samples from the bernbjerg formation (samples 1–3, fig. 9) all contain a pronounced archaean zircon age population with peak ages at 2.75–2.65 ga (comprising 7–10% in each sample), whereas only a single archaean zircon grain was found in the sample from the stratumbjerg formation (sample 4, fig. 9). the dominant age populations of the samples are present within the 2.0–1.6 ga interval (comprising 42–75% in each sample), although the relative proportions between the age populations vary. in the https://doi.org/10.34194/geusb.v55.8309 http://www.geusbulletin.org/ olivarius et al. 2023: geus bulletin 55. 8309. https://doi.org/10.34194/geusb.v55.8309 12 of 22 geusbulletin.org bernbjerg formation, an age population with peak age at 1.97 ga is dominant in two of the samples (samples 1 and 3, comprising 23–24%) and evident in the third (sample 2, comprising 10%), but not in the sample from the stratumbjerg formation. the dominant age population in the stratumbjerg formation sample has peak age at 1.88 ga (sample 4, comprising 39%), and this population is less pronounced and slightly older in the bernbjerg formation samples. an age population with peak ages at 1.65–1.63 ga is pronounced in the sample from stratumbjerg formation (sample 4, comprising 17%) and in two of the samples from the bernbjerg fig. 8 bioclasts and fracture fillings. a–b: recrystallised calpionellids filled with combinations of calcite, dolomite, ankerite and pyrite, and occasionally with internal porosity. calpionellids are present in bernbjerg and lindemans bugt formations. c–d: well-preserved calcispheres filled with calcite and small ankerite rhombs. calcispheres are present in palnatokes bjerg formation. e: opal precipitated in several zones in a fracture and succeeded by dolomite, which has replaced some opal along the contact (arrow). f: opal along fracture rims and succeeded by dolomite with lower fe content in the middle. fractures are most abundant in bernbjerg formation. see fig. 6 for abbreviations. a lindemans bugt fm, tl image, rø-1, 76.47 m b lindemans bugt fm, bse image, rø-1, 76.47 m c palnatokes bjerg fm, se image, rø-1, 13.31 m d palnatokes bjerg fm, bse image, rø-1, 13.31 m e bernbjerg fm, cn image, rø-1, 187.50 m f bernbjerg fm, bse image, rø-1, 187.50 m 100 μm 20 μm 10 μm5 μm 300 μm 500 μm d c pf dpf c a c o d d d o o https://doi.org/10.34194/geusb.v55.8309 http://www.geusbulletin.org/ olivarius et al. 2023: geus bulletin 55. 8309. https://doi.org/10.34194/geusb.v55.8309 13 of 22 geusbulletin.org formation (samples 2 and 3, both comprising 16%). a population with peak age of 1.74 ga is slightly less prominent in the same three samples (comprising 9–14% in each sample), whereas both age populations are less prominent in the last sample from the bernbjerg formation (sample 1; fig. 9). a wide range of mesoproterozoic zircon ages are found in all four samples (comprising 11–38% in each fig. 9 zircon u-pb age distributions of the upper jurassic – lower cretaceous sediments (1–4) compared to selected zircon ages from the region (a–g: strachan et al. 1995; watt et al. 2000; thrane 2002; elvevold et al. 2003; kalsbeek et al. 1993; leslie & nutman 2003; sláma et al. 2011; olivarius et al. 2018b; barham et al. 2020; olierook et al. 2020). the phanerozoic zircon ages in a–c are from the intruded caledonian granites. the sampling locations are shown in fig. 1. the ages are plotted using kernel density estimation (vermeesch 2012) and histograms with a bin interval of 25 million years. zircon ages with <10% discordancy are plotted. “n/n” denotes the number of concordant analyses out of the total number of analyses. see fig. 1 for locations. 0 30 15 0 5 10 pr ob ab ili ty pr ob ab ili ty pr ob ab ili ty pr ob ab ili ty pr ob ab ili ty n um be r o f g ra in s 0 3 6 0 8 4 0 8 4 0 30 15 0 20 10 pr ob ab ili ty pr ob ab ili ty pr ob ab ili ty 0 6 3 0 6 3 0 8 4 pr ob ab ili ty pr ob ab ili ty 4 3 2 1 f e d c b a pr ob ab ili ty g 0 40 20 age (ga)0.0 1.0 2.0 3.0 4.00.5 1.5 2.5 3.5 neoproterozoic mesoproterozoic palaeoproterozoic archaeanphanerozoic neoproterozoic metasediments eleonore bay supergroup nathorst land group middle–upper devonian sediments kap koltho� and kap graah groups upper jurassic sediments hall bredning group hareelv formation lower cretaceous sediments hold with hope group steensby bjerg formation upper jurassic sediments hall bredning group bernbjerg formation upper jurassic sediments hall bredning group bernbjerg formation upper jurassic sediments hall bredning group bernbjerg formation lower cretaceous sediments brorson halvø group stratumbjerg formation n/n = 83/115 this study n/n = 112/150 this study n/n = 103/135 this study n/n = 81/103 this study 2 samples, n/n = 120/120 sláma et al. 2011 6 samples, n/n = 660/945 olivarius et al. 2018b 5 samples, n/n = 524/524 sláma et al. 2011 6 samples, n/n = 490/960 olierook et al. 2020 archaean–palaeoproterozoic rocks crystalline basement complexes meso–neoproterozoic metasediments krummedal and smalle�ord sequences 6 samples, n/n = 114/181 strachan et al. 1995, watt et al. 2000, leslie & nutman 2003 15 samples, n/n = 127/147 kalsbeek et al. 1993, thrane 2002, elvevold et al. 2003 lower cretaceous sediments wollaston forland group lindemans bugt & palnatokes bjerg formations 7 samples, n/n = 459/935 barham et al. 2020 2.67 1.96 1.85 1.65 1.66 0.43 0.42 0.38 2.47 2.71 1.96 1.10 1.36 1.49 1.64 1.50 1.090.94 0.42 0.42 2.74 1.711.64 1.48 1.321.12 2.662.51 1.96 1.89 1.65 1.731.11 1.03 0.43 1.461.38 1.14 1.74 2.81 1.88 1.01 1.51 2.50 2.79 1.97 1.971.89 1.63 1.41 2.65 2.74 1.74 1.72 1.65 1.56 1.28 0.41 1.97 2.72 1.90 1.90 1.74 1.64 0.45 1.64 0.41 1.88 1.74 0.40 1.73 1.91 1.10 2.52 1.731.64 2.940.41 https://doi.org/10.34194/geusb.v55.8309 http://www.geusbulletin.org/ olivarius et al. 2023: geus bulletin 55. 8309. https://doi.org/10.34194/geusb.v55.8309 14 of 22 geusbulletin.org sample) with less pronounced age populations as compared to the palaeoproterozoic populations. few neoproterozoic zircons are encountered (comprising 1–6% in each sample), and an age gap occurs at 0.9–0.5 ga except for a few grains (fig.  9). small palaeozoic age populations of 2–4 grains are found in all samples with peak ages varying between 0.45 and 0.38 ga (comprising 2–5% in each sample). the oldest and youngest of these ordovician–devonian zircons, with ages of 474 ± 14 ma and 364 ± 3 ma, respectively, are both from the bernbjerg formation. 5. discussion 5.1 provenance analysis the mesoarchaean to palaeozoic zircon age populations found in the four analysed samples of upper jurassic – lower cretaceous sediments from wollaston forland are comparable (fig. 9). the age distributions of the studied sediments are all characterised by a dominance of 2.0–1.6 ga ages and by containing the full range of zircon ages within this time interval, although the relative proportion between the probabilities of the age populations varies. therefore, the sediments presumably have the same overall provenance in which the proportion between the palaeoproterozoic zircon ages varies. however, some of the differences in age distributions between the samples may possibly be the result of bias related to mineral separation with handpicking of grains for analysis (e.g. sláma & košler 2012; dröllner et  al. 2021). selected zircon age data from the literature are plotted in fig. 9 to facilitate comparison with possible sediment source rocks. as in the studied sediments, zircon age populations with a dominance of 2.0–1.6 ga ages occur in lower cretaceous and lower triassic sediments from northern hold with hope, c. 80 km south-south-west of the study area (fig.  1; fonneland et  al. 2004; sláma et  al. 2011). these sediments also contain archaean, mesoproterozoic and palaeozoic age populations with low probabilities, although the palaeozoic peak age is more pronounced than in the studied sediments (fig.  9). the age populations of upper jurassic sediments from southern jameson land c. 450 km south-south-west of the study area are similar overall to the studied sediments, but with significantly larger relative proportions of the archaean and mesoproterozoic age populations (olivarius et al. 2018b). lower jurassic sediments from southern jameson land have a significantly different age distribution; they exhibit a limited number of peak ages reflecting their local provenance from the liverpool land high that was elevated at the time (sláma et  al. 2011). the zircon age populations of carboniferous and devonian sediments along kong oscar fjord and inner kejser franz joseph fjord (moskusokse fjord) are comparable to those in the studied sediments, except that the mesoproterozoic and palaeozoic populations are more prominent in these older sediments; furthermore, archaean ages are more common in the carboniferous sediments (sláma et al. 2011). the neoproterozoic lyell land group and nathorst land group of the eleonore bay supergroup have age distributions that are distinctly different from each other (watt et al. 2000; dhuime et al. 2007; sláma et al. 2011; olierook et al. 2020). late mesoproterozoic zircons are dominant in the lyell land group along inner kong oscar fjord (segelsellskapet fjord) with subordinate early mesoproterozoic zircons, so these sediments show poor resemblance to the studied upper jurassic – lower cretaceous sediments. the nathorst land group is dominated by late palaeoproterozoic age populations and additionally contains several mesoproterozoic age populations in addition to a smaller archaean population; a sample of the intruded granites is included in fig.  9 to show their palaeozoic age (olierook et  al. 2020). the metasediments of the nathorst land group on wollaston forland are thus comparable to the studied sediments, except that the metasediments contain a higher proportion of mesoproterozoic zircons and no significant age populations in the 2.0–1.8 ga interval. the meso–neoproterozoic metasediments of the krummedal supracrustal sequence and smallefjord sequence have dominant zircon age populations of late palaeoproterozoic age and additionally contain mesoproterozoic, early neoproterozoic and palaeozoic populations (strachan et al. 1995; watt et al. 2000; leslie & nutman 2003). thus, they are lacking 2.0–1.8 ga age populations but otherwise resemble the age distributions of the studied sediments rather well. archaean and palaeoproterozoic crystalline basement complexes of the east greenland caledonides have age populations of 2.9–2.5 and 2.0–1.7 ga (kalsbeek et al. 1993; thrane 2002; elvevold et al. 2003). these are comparable to the oldest age populations found in the studied sediments, although the proportion of palaeoproterozoic ages relative to archaean ages is higher in the sediments than in the basement complexes. comparison between the samples analysed in this study and other sediments and possible sediment sources is made by multivariate statistical analysis by multidimensional scaling (mds) visualised in an mds diagram (fig.  10). here, similarities between samples are highlighted by solid lines revealing their proximity in kolmogorov-smirnov space, whereas dashed lines show smaller similarities. the sediments from this study have largest similarities to other upper jurassic and https://doi.org/10.34194/geusb.v55.8309 http://www.geusbulletin.org/ olivarius et al. 2023: geus bulletin 55. 8309. https://doi.org/10.34194/geusb.v55.8309 15 of 22 geusbulletin.org lower cretaceous sediments, though some of these are as far away as jameson land. the largest differences are found between the end members comprising the archaean–palaeoproterozoic rocks and the meso–neoproterozoic metasediments, indicating that a mixture of these is necessary to explain the range of ages encountered in the sediments. 5.2 sediment transport an overall north–south change in provenance in the upper palaeozoic to mesozoic succession in east greenland is evident in the concentration of palaeoproterozoic zircons in sediments on wollaston forland and hold with hope compared with sediments farther south, which contain a higher proportion of mesoproterozoic zircons (fig.  9). in particular, 2.0–1.8 ga zircons are abundant in the central and northern parts of east greenland in triassic, jurassic and cretaceous sediments (fonneland et al. 2004; sláma et al. 2011; this study). this implies that there must be pronounced differences between either age or extent, or both, of the various sediment source rocks in the northern versus southern parts of the caledonides. factors such as zircon fertility, sediment routing, recycling and methodological bias may also affect the provenance signal (e.g. dröllner et al. 2021). the primary source of the upper jurassic – lower cretaceous sediments on wollaston forland comprised the crystalline rocks of the east greenland caledonides or their derived sediments, or both. the archaean and palaeoproterozoic ages of the basement in payer land (elvevold et al. 2003) match the oldest age populations of the studied sediments, so they may have been supplied from the crystalline complexes present west of wollaston forland (fig.  1). an additional sediment source is necessary to account for the late palaeoproterozoic peak age of 1.65–1.63 ga and the range of mesoproterozoic ages present in the studied sediments (fig. 9), so this input must originate from erosion of meso-neoproterozoic metasediments or palaeozoic sediments, or both. the peak zircon age of 1.66–1.63 ga is evident and often dominant in the sediments and metasediments of the east greenland caledonides (fig.  9), except for the lyell land group in which the peak ages are restricted to 1.5 and 1.1 ga (sláma et al. 2011). peak ages of 1.5 and 1.1 ga are also present in the krummedal supracrustal sequence, the nathorst land group and younger sediments such as the devonian and carboniferous, but are not pronounced in the studied sediments. however, these peak ages are distinct in the upper jurassic sediments on jameson land that are age-equivalent to the bernbjerg formation (fig.  9). this is due to the lower proportion of mesoproterozoic zircon ages in the northern part of the east greenland caledonides and in the studied sediments, which makes the individual age populations in this interval less distinct. some change in the drainage pattern must have occurred between the kimmeridgian – early volgian and the albian, since the distinct  archaean and 1.97 ga populations in the bern bjerg formation are not evident in the stratumbjerg formation where a 1.88 ga population is dominant, as observed in aptian sediments on northern hold with hope. ages corresponding to these populations are present in different crystalline basement complexes (fig. 9). zircon age distributions of the bernbjerg, lindemans bugt, palnatokes bjerg and stratumbjerg formations from western wollaston forland are comparable to each other (barham et al. 2020) and broadly comparable to the new results from northern wollaston forland (fig. 9). however, the pronounced peak age of 1.97 ga in the three samples from the bernbjerg formation in northern wollaston forland is not present in samples from the west. likewise, caledonian zircons are also virtually absent in the west in contrast to the northern part of wollaston forland. thus, although much of the sediment on northern wollaston forland has been produced from reworking of sediment from western wollaston forland (fig. 11), there must also have been an additional source that supplied sediment to the half-graben in which the rødryggen-1 and brorson halvø-1 boreholes are situated. this additional sediment was probably supplied by axial transport from the north in accordance with the general depositional pattern in the late jurassic (surlyk 2003). this is compatible with the abundance of palaeoproterozoic basement to the north fig. 10 multidimensional scaling (mds) diagram of zircon u-pb age data. plotted using kolmogorov-smirnov (k-s) dissimilarity (vermeesch et al. 2016). the nearest neighbours in k-s space are shown by solid lines and the second nearest by dashed lines. see fig. 10 for sample information. −0.6 −0.4 −0.2 0.0 0.2 0.4 0.6 –0 .6 –0 .4 –0 .2 0. 0 0. 2 0. 4 0. 6 dim 1 d im 2 a b c d e f g 1 2 3 4 https://doi.org/10.34194/geusb.v55.8309 http://www.geusbulletin.org/ olivarius et al. 2023: geus bulletin 55. 8309. https://doi.org/10.34194/geusb.v55.8309 16 of 22 geusbulletin.org and the presence of caledonian granites in this area, though it is unknown if these granites were exposed in the mesozoic. changes in sediment composition are also evident in the clay mineral composition where chlorite and vermiculite were added upwards in the succession (figs 3, 4). another difference between the bernbjerg formation and the younger deposits is evident in the zircon age distributions from northern wollaston forland where the stratumbjerg formation lacks the archaean and 1.97 ga age populations that are present in the bernbjerg formation (fig.  9). thus, the rifting activity probably changed the erosional pattern resulting in altered sediment transport pathways. the kaolinite that is overgrown by authigenic quartz in the mudstones must comprise detrital kaolinite since they do not occur in well-defined booklets (fig. 7c). the climate was humid subtropical at the time of deposition (surlyk 2003), so kaolinite is the most likely clay mineral to have formed in the hinterland (rateev et al. 2008), which explains its abundance in the mudstones. the presence of detrital kaolinite with preserved euhedral shapes points to a short transport distance from source area to place of deposition. this agrees with the interpreted proximity to the deltaic coast of north-east greenland (hovikoski et al. 2023b) with high mountains consisting of readily erodible material in the immediate hinterland (henriksen & higgins 2008). 5.3 tectonic regime comparing the mineralogy of the carbonate-cemented mudstones with uncemented mudstones shows that the mudstones had similar initial mineralogical compositions, with the exception of the carbonate bioclasts, which represent the only additional component occurring in the cemented intervals (figs 3, 4). thus, in general, sediments with similar composition were supplied to the basin during deposition of the studied stratigraphic units, although the proportion between grains and clay minerals varies in relation to the grain size (figs 3, 4). other differences in the detrital mineralogy include the varying content of apatite clasts and the addition of vermiculite and chlorite to the clay mineral assemblages upwards in the succession. the variations in detrital mineralogy and grain size of the studied sediments are the results of (1) changing tectonic regime of the depositional setting that changed the sediment transport pathways and sea-bottom topography, and (2) rotational block faulting that caused increasing basin depth and sediment starvation up through the succession (hovikoski et  al. 2023a, this volume). rifting also influenced the amount of deoxygenation and thereby the diagenetic evolution of the sediments as reflected in the varying amount of precipitated pyrite. the calcareous bioclast abundance is largest in transgressive intervals and in some condensed intervals as is also the case for apatite clasts. the marine shelf setting of the bernbjerg formation with sediment supply from the deltaic coastline to the west (fig. 11a) is reflected in the sandy component of the fig. 11 inferred structural setting of northern wollaston forland in late jurassic to early cretaceous time when submarine deposition of mudstones took place during (a) early rifting, (b) rifting climax and (c) waning rifting in west (not waning in east, i.e. falskebugt member). locations of the rødryggen-1 (rø-1) and brorson halvø-1 (bh-1) boreholes are shown as well as sediment transport directions. a) early rift: bernbjerg formation b) rift climax: lindemans bugt formation c) waning rift: palnatokes bjerg formation & stratumbjerg formation bh-1rø-1 n bh-1rø-1 n bh-1rø-1 n axial sediment transport basinwards sediment transport https://doi.org/10.34194/geusb.v55.8309 http://www.geusbulletin.org/ olivarius et al. 2023: geus bulletin 55. 8309. https://doi.org/10.34194/geusb.v55.8309 17 of 22 geusbulletin.org mudstones with storm-wave influence in the lower part of the cored succession and in the presence of abundant silt and fine sand-sized quartz grains (fig. 6c). faulting intensified (surlyk 2003), linked to the onset of coarse marine sedimentation proximally (lindemans bugt formation). at the culmination of rotational block faulting, a fault was created west of the study area resulting in westwards tilting of the resulting basin (fig.  11b; hovikoski et al. 2023b). the submarine fault scarp west of this half-graben blocked sediment input from the mainland, but some sediment was supplied into the basin by gravity flows originating at the fault scarp. the fairly short distance to the fault scarp resulted in the input of some silt-sized detrital grains despite the relatively sediment-starved environment. clasts of quartz, feldspar, mica and apatite (fig. 6e) were supplied by mass flows in addition to the detrital clay minerals. vermiculite is absent in the bernbjerg formation in the rødryggen-1 core but is found in all samples of the lindemans bugt and palnatokes bjerg formations (fig. 3). the cause of this change is not clear, but it may reflect the changed setting caused by block faulting whereby sediment supplied to the rødryggen-1 drill site was sourced by the submarine fault scarp to the west during rift climax, and this input mixed with sediment supplied by axial transport during the waning rift phase (fig. 11c). the high gamma-ray values of the lindemans bugt formation in the rødryggen-1 borehole (fig. 3) are compatible with the fine-grained nature of the deposits in the centre of the half-graben where fine-crystalline vermiculite is dominant and coarse-crystalline kaolinite becomes less abundant, whereas the formation is only thinly preserved farther to the east in the brorson halvø-1 borehole (fig. 4). vermiculite does not occur in the succession in the brorson halvø-1 core, except for minor amounts in the uppermost sample from stratumbjerg formation. this is compatible with the larger distance to the submarine fault crest to the west that primarily fed vermiculite to the proximal part of the half-graben. the palnatokes bjerg formation is more sediment-starved than the lindemans bugt formation due to transgression and waning of active rifting (fig.  11c; surlyk 2003). this is reflected in the high calcareous content suggestive of slow clastic deposition. the presence of occasional muddy sandstones, however, implies that gravity-flow processes were still operative. in the brorson halvø-1 core, chlorite occurs in the rift climax and waning rift samples, whereas it is only found in waning rift samples in the rødryggen-1 core, and in smaller amounts (figs 3, 4). this indicates that chlorite was supplied from the north-north-west by axial transport and deposited primarily in the distal part of the half-graben since the proximal part was mainly fed by the submarine fault scarp to the west. the abundance of chlorite may have resulted from erosion of the meso–neoproterozoic metamorphic rocks north-west of wollaston forland that were exposed due to rift faulting, since chlorite often originates from such lithologies (nielsen et  al. 2015). the provenance signature of the stratumbjerg formation near the brorson halvø-1 drill site is also indicative of a change in source area since the archaean zircon age population is absent, in contrast to the bernbjerg formation (fig. 9), which signifies a change in sediment source from the crystalline rocks to the west to the metamorphic and crystalline rocks to the north-west of the half-graben. this change is not evident in the zircon age distributions of the palnatokes bjerg and lindemans bugt formations reported by barham et al. (2020) because their samples were taken from localities within the proximal half graben that were linked directly to the coastline. 5.4 diagenetic evolution textural relationships in sediments of the rødryggen-1 and brorson halvø-1 cores have been used to determine the diagenetic sequence. a diagenetic process scheme is established to highlight the relative importance of each process (fig. 12), as discussed next. the sedimentary succession in the cores has poor reservoir quality since the few sandstone intervals are thin and muddy. the reservoir properties are poorest in the intervals with pervasive carbonate cementation. the formation of secondary porosity by partial dissolution of bioclasts and feldspars and by fracture formation has only had a minor influence on the total porosity. the permeability is only slightly affected by the dissolution process since it was restricted to local clasts. the open fractures, however, have presumably increased the permeability significantly. 5.4.1 eogenetic processes pyrite framboids precipitated early in the sediments in association with bacterial sulphate reduction of organic matter (fig.  7a). the membrane of organic matter in shells probably promoted early pyrite formation within many of the bioclasts (fig. 8a). euhedral pyrite formed only during the rift climax, corresponding primarily to the lindemans bugt formation where the largest amounts of pyrite are found (figs 3, 4, 5c, 7b). the small euhedral crystals suggest syngenetic formation of pyrite indicating that the chemocline moved above the sediment–water interface and anoxic conditions may have prevailed during the rift climax phase (tribovillard et al. 2006) or the crystals may have formed later diagenetically. dysoxic conditions were dominant when the rifting was still at an initial stage during deposition of the bernbjerg formation, and when rifting was waning as seen by the increased bioturbation and low https://doi.org/10.34194/geusb.v55.8309 http://www.geusbulletin.org/ olivarius et al. 2023: geus bulletin 55. 8309. https://doi.org/10.34194/geusb.v55.8309 18 of 22 geusbulletin.org pyrite content of the palnatokes bjerg and stratumbjerg formations. during sulphidic bottom-water conditions, the reactivity and abundance of fe minerals control the amount of pyrite that can form, whereas pyrite formation is controlled by the reactivity and abundance of organic matter during oxygenated bottom-water conditions (berner 1985). thus, the low amounts (0–1%) of total organic carbon (toc) present in the palnatokes bjerg and stratumbjerg formations largely precluded pyrite precipitation, whereas higher toc (2–6%) in the bernbjerg and lindemans bugt formations (bojesen-koefoed et  al. 2023b, this volume) will have favoured this process. the calcite and dolomite cements probably formed shortly after deposition. the sparry carbonate cement within calpionellids in bernbjerg and lindemans bugt formations must have formed early after deposition since the calpionellids have not been deformed by mechanical compaction (fig. 8a, b). the presence of bioclasts is important for the precipitation of dolomite and ankerite (e.g. hendry et al. 2000; burns et al. 2005) since they act both as nucleation sites for the crystals and as a source of carbonate. the palnatokes bjerg formation was sediment-starved so pelagic lime mud formed a significant component of the sediment that accumulated on the sea floor, and the pelagic carbonate resulted in micritic calcite matrix (fig. 7f). in the bernbjerg and lindemans bugt formations, the laminated mudstones signify reducing conditions during deposition as also reflected in the relatively high content of organic matter. in the palnatokes bjerg formation, the bioturbation intensity reveals more oxygenated depositional conditions, and many of the bioclasts were preserved in this formation (figs 8c, d). partial dissolution of calpionellids and other bioclasts in the bernbjerg and lindemans bugt formations resulted in carbonate-cemented layers with a large bioclast content, caused by a primary heterogeneity (fig.  7e). these layers possibly represent bioclast concentrations produced by reworking at discrete flooding surfaces (e.g. burns et  al. 2005). this is also suggested by the cemented mudstones that typically occur at the top of a few metres of upward-coarsening successions or in the finest grained intervals at the base of upward-coarsening cycles. the fewer cemented intervals in the lindemans bugt formation in comparison to the bernbjerg formation (fig. 3) are thus in accordance with the progressive deepening of the setting, where the influence of minor relative sea-level fluctuations decreased steadily. 5.4.2 mesogenetic processes the presence of authigenic mixed layer illite-smectite and illite in the studied mudstones indicates that illitisation of smectite has occurred, thereby providing silica and cations for other mineral reactions such as the formation of quartz and ankerite (fig.  12). significant illitisation may, however, be contradicted by the presence of k-feldspar in the investigated mudstones, as k-feldspar disappears in shales from most wells below 2.5 km in the northern north sea due to the illitisation fig. 12 diagenetic process scheme of the authigenic changes that have occurred in the rødryggen-1 and brorson halvø-1 successions. see section 5.4 for explanation. diagenetic process sulphate reduction bioclast alteration smectite illitization carbonate transformation fracture formation meteoric water �ushing diagenetic regime eogenesis eogenesis mesogenesis mesogenesis telogenesis telogenesis temperature depositional depositional >80°c 80–100°c <65°c <50°c dissolution organic matter, fe-minerals bioclasts, carbonate ooze smectite, k-feldspar dolomite, smectite bulk mudstone feldspar, mica precipitation pyrite calcite, dolomite illite, quartz ankerite, barite opal, dolomite kaolinite porosity decreased decreased decreased decreased increased slightly increased slightly permeability decreased decreased decreased decreased increased decreased slightly requirements s from seawater mg from seawater detrital smectite ba from initial seawater source of water supply of meteoric water morphology https://doi.org/10.34194/geusb.v55.8309 http://www.geusbulletin.org/ olivarius et al. 2023: geus bulletin 55. 8309. https://doi.org/10.34194/geusb.v55.8309 19 of 22 geusbulletin.org of smectite (pearson & small 1988). if the deposited clay minerals were rich in illite, however, the pore fluids may not have been undersaturated with respect to potassium and consequently less aggressive towards k-feldspar during early diagenesis. hence, k-feldspar may have survived until deep burial. the iron necessary to form ankerite may have been provided by smectite illitisation and it may have formed at the expense of dolomite cement. barite is typically a late authigenic phase because the pore fluids need to be very concentrated before they contain sufficient ba for barite formation. as expected, barite formed late in the studied sediments as seen by the crystal habit of barite as euhedral overgrowths (fig. 7d), and the string-like occurrence of the precipitated barite identifies the transport route of the last pore fluid. the diagenetic alteration of the mudstones is indicative of the maximum burial depths to which they have been exposed prior to structural inversion. the presence of quartz overgrowths shows that the sediments must have been exposed to temperatures of at least c. 80°c that is necessary for the growth onset and less than c. 100°c since they are thin and scattered (bjørlykke & jahren 2015). precipitation of quartz in mudstones can be sourced by the smectite to illite transformation, which generally occurs at temperatures of 60–100°c (thyberg et al. 2009) and is likely to have happened in the studied sediments. barite precipitated after quartz and ankerite in the mudstones as the last mineral phase during deep burial, a process that has been found to occur at temperatures of 83–105°c in sandstones (burley et  al. 1989). thus, the precipitation temperatures of the authigenic phases that formed at deepest burial correspond to maximum burial depths of c. 2.1–2.6 km assuming a surface temperature of 20°c and a palaeogeothermal gradient of 30°c/km (japsen et  al. 2021). this is less than the estimated uplift of c. 2.8 km based on thermochronological data from upper jurassic sediments from jameson land c. 450 km towards south-south-west, in which the diagenesis has progressed further (green & japsen 2018; olivarius et al. 2018a). 5.4.3 telogenetic processes the fracturing that has occurred in the sediments must have happened late during uplift since the thin fractures cross-cut the authigenic phases in the sediments. fracturing is presumed to have accompanied exhumation and pressure release since cooling of the sediments was necessary to decrease their elasticity enough for fracturing (e.g. gale et al. 2014). the fractures formed mainly in the cemented intervals because they had least elasticity. the late timing of the fracturing is also testified by the opal infill, which requires low temperatures of <65°c for precipitation (weibel et al. 2010). the pore fluids must have been oversaturated with silica in the beginning since opal often formed along the rim of the fractures (fig. 8e). formation of fe-rich dolomite then took over in accordance with the expected composition of the pore fluid since the fractured mudstones are cemented with dolomite and ankerite. the pore fluid became depleted in fe, so the last dolomite that precipitated was fe-poor (fig.  8f). the last fracturing episode happened so late during exhumation that no mineral phases were precipitated in the fractures, which thus contribute minor secondary porosity. however, it is difficult to differentiate natural open fractures, formed in the subsurface, from artefacts produced during drilling and drying of the core. introduction of meteoric water is a common mechanism for kaolinite formation (bjørlykke 1998), as this process could not take place after deposition in the marine environment. the partially dissolved feldspar grains have not been deformed so the secondary porosity has been preserved (fig. 6d), which indicates that the dissolution and kaolinisation of detrital phases may have happened late during the exhumation (fig. 12). the fracturing that has occurred during late uplift may have caused a flow of meteoric water through the sediment that was sufficient for kaolinite formation. this may also explain why the latest generation of fractures is not cemented since the pore fluids had a low saturation. 5.5 implications for sediment composition in the norwegian–greenland seaway the presence of archaean zircon grains in the norwegian sea is often considered diagnostic of sediment supply from east greenland since archaean zircons are scarce in sediment eroded off the fennoscandian shield (e.g. morton et al. 2008). only a single archaean zircon grain was retrieved from the sample from the stratumbjerg formation on wollaston forland (fig. 9). however, the absence or scarcity of archaean zircons cannot be considered unambiguous proof of a fennoscandian source in the offshore sediments. this is particularly the case in sediment derived from the central or northern part of the east greenland caledonides where palaeoproterozoic basement is most abundant, whereas archaean basement is more abundant in the southern caledonides (thrane 2002). this is reflected in the geographical differences evident in the zircon age distributions of post-caledonian sediments in east greenland (sláma et al. 2011; olivarius et al. 2018b). in sediments with an east greenland source, the present results highlight how a high proportion of mesoproterozoic and latest palaeoproterozoic zircon ages testifies to a provenance from the southern east greenland caledonides, whereas a higher proportion of late palaeoproterozoic https://doi.org/10.34194/geusb.v55.8309 http://www.geusbulletin.org/ olivarius et al. 2023: geus bulletin 55. 8309. https://doi.org/10.34194/geusb.v55.8309 20 of 22 geusbulletin.org (2.0–1.7 ga) ages is indicative of a more northern sediment source. this is reflected in the distribution of zircon age populations in sediments in the western part of the norwegian sea such as in the upper cretaceous – paleocene succession (fonneland et al. 2004; morton et al. 2005). time-equivalent sediments in the norwegian sea have presumably experienced some of the same diagenetic reactions as the studied sediments, specifically, initial sulphate reduction causing pyrite precipitation, eogenetic bioclast alteration causing calcite-dolomite cementation, mesogenetic smectite illitisation and quartz precipitation and carbonate transformation into ankerite (fig. 12). 6. conclusions this study illustrates that changes in mudstone composition can be induced by half-graben evolution and that the composition may also vary in relation to the position in the rift basin and the sea-bottom topography. this knowledge can be applied to predict variations in sediment composition in underexplored half-graben settings. the diagenetic evolution includes processes related to the different diagenetic regimes that the mudstones have experienced. during early diagenesis, sulphate reduction caused pyrite formation, and bioclast alteration resulted in precipitation of calcite and dolomite. during burial diagenesis, illite and quartz formed due to smectite illitisation, and ankerite and barite precipitated because of carbonate transformation. during uplift, opal and dolomite precipitated in the earliest fractures, and kaolinite formed due to meteoric water flushing. the provenance analysis of sand-rich intervals shows that the zircon age patterns of the studied sediments are most similar to other upper jurassic – lower cretaceous sandstones from east and north-east greenland. this is revealed by mds where it is evident that the archaean–palaeoproterozoic crystalline basement complexes and the meso–neoproterozoic metasediments comprise two end members, so their derived detritus must have been mixed to produce the mesozoic sediments. acknowledgements technical assistance by helene almind, kirsten fries, john boserup, fiorella f. aguilera, mojagan alaei, michael s. nielsen and jette halskov is much appreciated. valuable advice was provided by emma sheldon, jon ineson, henrik vosgerau, holger lindgreen and tonci balic-zunic. the authors would like to thank the reviewers chris kirkland and kevin taylor for insightful comments that helped improve the manuscript. additional information funding statement funding for drilling of the rødryggen-1 and brorson halvø-1 boreholes and studies of the cores was provided by a consortium of oil companies and the geological survey of denmark and greenland (geus). author contributions mo: wrote the manuscript in cooperation with the other authors. mo, ak, rw: performed the mineralogical and petrographic work. tk: processed the radiometric analyses. jh: carried out the sedimentological work. competing interests none. additional files: the full table of zircon u-pb data is provided as supplementary file s1, available at https://doi.org/10.22008/fk2/kel0a6. references alsen, p., piasecki, s., nøhr-hansen, h., pauly, s., sheldon, e. & hovikoski, j. 2023: stratigraphy of the upper jurassic to lowermost cretaceous in the rødryggen-1 and brorson halvø-1 boreholes, wollaston forland, north-east greenland. geus bulletin 55, 8342 (this volume). https:// doi.org/10.34194/geusb.v55.8342 barham, m., kirkland, c.l., hovikoski, j., alsen, p., hollis, j. & tyrrell, s. 2020: reduce or recycle? revealing source to sink links through integrated zircon–feldspar provenance fingerprinting. sedimentology 68, 531–556. https://doi.org/10.1111/sed.12790 berner, r.a. 1985: sulphate reduction, organic 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https://doi.org/10.34194/geusb.v1.4674 https://doi.org/10.1016/j.marpetgeo.2013.03.009 https://doi.org/10.1016/j.marpetgeo.2013.03.009 https://doi.org/10.34194/geusb.v46.6521 https://doi.org/10.1016/s0301-9268(01)00198-x https://doi.org/10.1016/s0301-9268(01)00198-x https://doi.org/10.3997/1365-2397.2009003 https://doi.org/10.1016/j.chemgeo.2006.02.012 https://doi.org/10.1016/j.chemgeo.2006.02.012 https://doi.org/10.1016/j.chemgeo.2012.04.021 https://doi.org/10.1016/j.chemgeo.2012.04.021 https://doi.org/10.1016/j.sedgeo.2016.01.009 https://doi.org/10.1016/j.sedgeo.2016.01.009 https://doi.org/10.1144/jgs.157.5.1031 https://doi.org/10.1016/j.sedgeo.2010.04.008 https://doi.org/10.1016/j.sedgeo.2010.04.008 mudstone diagenesis and sandstone provenance in an upper jurassic – lower cretaceous evolving half-graben system, wollaston forland, north-east greenland 1. introduction 2. geological setting 3. methodology 3.1 petrography 3.2 zircon u-pb geochronology 4. results 4.1 lithology 4.2 detrital components 4.3 authigenic minerals 4.4 bioclasts 4.5 fractures 4.6 zircon u-pb ages 5. discussion 5.1 provenance analysis 5.2 sediment transport 5.3 tectonic regime 5.4 diagenetic evolution 5.4.1 eogenetic processes 5.4.2 mesogenetic processes 5.4.3 telogenetic processes 5.5 implications for sediment composition in the norwegian-greenland seaway 6. conclusions acknowledgements additional information references figures fig. 1 geological map of north-east greenland based on stemmerik et al. (1997), henriksen et al. (2008) and kalsbeek et al. (2008a). detailed map of wollaston fig. 2 stratigraphic scheme with vertical lines showing the rødryggen-1 (rø-1) and brorson halvø-1 (bh-1) cored successions. al: albrechts bugt mb. b: bernbjerg fm. ba: bastians dal fm. j: jakobsstigen fm. l: laugeites ravine mb. li: lindemans bugt formation. li (s): lindemans bugt fm (storsletten mb). mu: muslingebjerg fm. n: niesen mb. pa: palnatokes fig. 3 mineralogy from xrd of the rødryggen-1 core plotted with the sedimentological log. chronostratigraphy from alsen et al. (2023, this volume). see fig. 4 for legend. chronostrat.: chronostratigraphy. lithostrat.: lithstratigraphy. gr: gamma ray. api: american petroleum units. fig. 4 mineralogy from xrd of the brorson halvø-1 core shown alongside the sedimentological log. chronostratigraphy from alsen et al. (2023, this volume). chronostrat.: chronostratigraphy. lithostrat.: lithstratigraphy. gr: gamma ray. api: american petroleum units. fig. 5 core photos of the mudstone texture of cemented versus uncemented mudstones. bh-1: brorson halvø-1 core. rø-1: rødryggen-1 core. a: laminated sandy mudstone. b: cemented mudstone. c: laminated pyritic mudstone. d: bioturbated mudstone. e: bioclastic hematitic cemented mudstone. f: bioturbated mudstone. fig. 6 texture and detrital phases. a: typical mudstone texture. b: cemented mudstone texture. c: quartz occurs as siltand sand-sized grains in the mudstones. d: partially dissolved albite grain. e: apatite occurs as detrital clasts. f: most clay minerals are detrital such as chlorite. fig. 7 authigenic minerals. a: framboidal pyrite has often formed in connection with organic matter. b: euhedral pyrite has only formed in lindemans bugt formation where it has overgrown pyrite framboids. c: quartz has overgrown detrital kaolinite. d: barite has overgrown ankerite rhombs. e: dolomite is the dominant carbonate cement in bernbjerg and lindemans bugt formations. f: calcite is the dominant carbonate cement in palnatokes bjerg formation. see fig. 6 for abbreviations. fig. 8 bioclasts and fracture fillings. a–b: recrystallised calpionellids filled with combinations of calcite, dolomite, ankerite and pyrite, and occasionally with internal porosity. calpionellids are present in bernbjerg and lindemans bugt formations. c–d: well-preserved calcispheres filled with calcite and small ankerite rhombs. calcispheres are present in palnatokes bjerg formation. e: opal precipitated in several zones in a fracture and succeeded by dolomite, which has replaced some opal along the contact (arrow). f: opal along fracture rims and succeeded by dolomite with lower fe content in the middle. fractures are most abundant in bernbjerg formation. see fig. 6 for abbreviations. fig. 9 zircon u-pb age distributions of the upper jurassic – lower cretaceous sediments (1–4) compared to selected zircon ages from the region (a–g: strachan et al. 1995; watt et al. 2000; thrane 2002; elvevold et al. 2003; kalsbeek et al. 1993; leslie & nutman 2003; sláma et al. 2011; olivarius et al. 2018b; barham et al. 2020; olierook et al. 2020). the phanerozoic zircon ages in a–c are from the intruded caledonian granites. the sampling locations are shown in fig. 1. the ages are plotted using kernel density estimation (vermeesch 2012) and histograms with a bin interval of 25 million years. zircon ages with <10% discordancy are plotted. “n/n” denotes the number of concordant analyses out of the total number of analyses. see fig. 1 for locations. fig. 10 multidimensional scaling (mds) diagram of zircon u-pb age data. plotted using kolmogorov-smirnov (k-s) dissimilarity (vermeesch et al. 2016). the nearest neighbours in k-s space are shown by solid lines and the second nearest by dashed lines. see fig. 10 for sample information. fig. 11 inferred structural setting of northern wollaston forland in late jurassic to early cretaceous time when submarine deposition of mudstones fig. 12 diagenetic process scheme of the authigenic changes that have occurred in the rødryggen-1 and brorson halvø-1 successions. see section 5.4 for explanation. geological survey of denmark and greenland bulletin 26, 2012, 73-76 73 ablation observations for 2008–2011 from the programme for monitoring of the greenland ice sheet (promice) robert s. fausto, dirk van as and the promice project team* recent estimates from the glaciological community agree that the greenland ice sheet is losing mass at an accelerated pace due to climate change (velicogna 2009; khan et al. 2010; rignot et al. 2011). this has caught the attention of the public and policy makers due to the potential impact on sea-level rise (dahl-jensen et al. 2009). the mass loss can be attributed approximately equally to increases in meltwater runoff from surface melt and iceberg production (van den broeke et al. 2009). the robustness of mass-balance predictions relies heavily on observational data from the greenland ice sheet and in recent years the need for frequent, reliable surface mass-balance measurements has increased (ipcc 2007; dahl-jensen et al. 2009). in anticipation of this need, the programme for monitoring of the greenland ice sheet (promice) was initiated in 2007, delivering in situ data from a network of automatic weather stations (aws) covering eight different regions of the ice sheet (fig. 1; van as et al. 2011). apart from the direct insight into the surface mass balance provided by these stations, the in situ data are also valuable for calibrating and validating melt estimates from remote sensors and surface mass-balance models (dahl-jensen et al. 2009). in this paper, we present the ablation records for the promice awss for 2008–2011, and the impact of the extraordinary atmospheric conditions on ablation in 2010 (tedesco et al. 2011) are compared to the other years. promice automatic weather stations the promice network in greenland currently consists of eight sites with two (or three) awss placed at different elevations (fig. 1; table 1; ahlstrøm et al. 2008) with a total of 18 stations. at each site, one station is typically located in the lower ablation zone close to the margin and the other in the upper ablation zone. exceptions are station kan_u, which is placed in the lower part of the accumulation zone, and station tas_u, which is placed well below the equilibrium-line altitude for reasons of accessibility. each station measures all relevant meteorological parameters and ice and snow ablation (fig. 2). it measures and stores data every ten minutes * andreas p. ahlstrøm, signe b. andersen, morten l. andersen, michele citterio, karen edelvang, signe h. larsen, horst machguth, søren nielsen and anker weidick. 80°n 75°n 70°n 65°n 60°n 80°w 60°w 40°w 50°w 40°w 20°w 20°w 0°w 2250 3000 2750 2500 1000 1500 1750 20 0022 50 27 5025 00 22 50 17 50 1500 1 2 5 0 2000 greenland 500 km qas nuk upe thu kpc sco tas kan fig. 1. map of greenland showing the locations of the promice automatic weather stations. each dot represents two or three stations. station names are found in table 1. dashed lines: elevation contours. © 2012 geus. geological survey of denmark and greenland bulletin 26, 73–76. open access: www.geus.dk/publications/bull 7474 with the exception of the wind-speed observations that give the mean wind speed since the last measurement cycle, and the gps measurements, which follow the transmission schedule. during winter (day 301 to day 99 of the year) the acquired data are transmitted once a day at midnight to limit power consumption when solar power is not available. during summer (day 100 to day 300) the data are transmitted hourly. the transmissions consist of daily or hourly average values of the more variable quantities such as temperature or radiation. values of less variable quantities, such as surface height and station tilt are appended once every six hours in summer and once a day in winter for all daily transmissions. instrumentation and accuracy the promice awss are equipped with an ørum & jensen nt1400 pressure transducer assembly and two campbell scientific sr50a sonic rangers that monitor surface-height change caused by accumulation and ablation. the pressure transducer assembly consists of a (non-freezing) liquid-filled hose with a pressure transducer located at its end/bottom. the hose is drilled into the ice. the pressure signal registered by the transducer is that of the vertical liquid column over the sensor, which can be scaled to depth using the density of the liquid. the free-standing aws tripod (fig. 2) moves down with the ablating surface and the hose melts out of the ice, reducing the hydrostatic pressure from the vertical liquid column in the hose. the reduction in pressure provides the ablation totals. the assembly was first constructed and implemented by bøggild et al. (2004). since the first successful tests in 2001, we have developed the system and now use a larger diameter hose and have changed from relative to absolute pressure transducers. we also use an ethylene-glycol solution, instead of alcohol and secure the upper reference level of the assembly (in the form of a bladder) to the mast instead of leaving it exposed on the ice surface. by measuring on a (sub-)daily timescale, the pressure transducer assembly is well suited to monitor ice ablation in remote regions, with clear advantages over other well-established methods. for instance, the stake readings providing information about surface-height change are done at every maintenance visit, so the records are limited to the visits. the accuracy of the sr50a sonic ranger is given by the manufacturer (campbell scientific) to be ±1 cm or ±0.4% of the measuring height after temperature correction. this was confirmed over 2.3 months of a virtually accumulation-free wintertime period at sco_u, during which time standard deviations of 1.7 cm and 0.6 cm were found (after spike removal), corresponding to 0.7% and 0.6% of the distance measured by the two sonic rangers, respectively. however, the precision of the readings from these sensors may reduce over time as the sensors degrade due to continuous cycles of moisture freezing on and melting off them. a major problem with sonic rangers in surface-mass balance studies is that they need to be mounted on the stake assemblies drilled into station latitude longitude elevation start date name (°n) (°w) (m a.s.l.) kpc_l* 79°55´ 24°05´ 380 17 july 2008 kpc_u 79°50´ 25°10´ 870 17 july 2008 sco_l 72°14´ 26°49´ 470 21 july 2008 sco_u 72°24´ 27°15´ 1000 21 july 2008 tas_l 65°38´ 38°54´ 270 23 august 2007 tas_u 65°42´ 38°52´ 580 15 august 2007 qas_l 61°02´ 46°51´ 310 24 august 2007 qas_u 61°11´ 46°49´ 890 7 august 2008 nuk_l 64°29´ 49°32´ 560 20 august 2007 nuk_u 64°30´ 49°16´ 1140 20 august 2007 nuk_n 64°57´ 49°53´ 930 26 july 2010 kan_l 67°60´ 50°70´ 670 1 september 2008 kan_m 67°40´ 48°49´ 1280 1 september 2008 kan_u 67°00´ 47°10´ 1830 4 april 2009 upe_l 72°54´ 54°18´ 230 17 august 2009 upe_u 72°53´ 53°32´ 980 17 august 2009 thu_l 76°24´ 68°16´ 570 9 august 2010 thu_u 76°25´ 68°09´ 770 9 august 2010 table 1. promice automatic weather station metadata (status 2011) *l: lower station, m: middle station, u: upper station, n: new station. 1 2 3 4 5 7 8 5 9 7 6 10 11 fig. 2. the promice automatic weather station upe_l photographed on 17 august 2009. 1: radiometer. 2: inclinometer. 3: satellite antenna. 4: anemometer. 5: sonic rangers. 6: thermometer and hygrometer. 7: pressure transducer. 8: solar panel. 9: data logger, barometer and gps. 10: battery box with 4 × 28 ah batteries. 11: 8-level thermistor string. 75 the ice. during a single melt season, these stake assemblies can melt out several metres, often causing them to move or even collapse during strong winds. in comparison, the pressure transducer assembly is operational until it has melted out of the ice, which can take several years depending on drill depth and the local ablation rate. this reduces the need for annual station visits, and thereby the considerable expenses associated with logistics in greenland. the measurement uncertainty of the pressure transducer sensor given by the manufacturer is 2.5 cm. the mean standard deviation of the upe and sco pressure transducer readings outside the ablation season is found to be ±1 cm, indicating a small random error comparable to that of the sonic ranger. the precision of the pressure transducer readings falls over time as the transducer degrades due to the continuous pressure on the sensor. sensitivity drift defines the amount by which an instrument’s sensitivity varies as ambient conditions change. calibration tests of the pressure transducer show that the sensor sensitivity drift amounts to 1.6% on average for a four-year measuring period (0.4% per year), suggesting that drift is not a large source of error. the pressure transducer is especially suitable for high ablation areas of >3 m per year because it is independent of the stakes drilled into the ice (fig. 2). ablation records for 2008–2011 the promice ablation records presented in table 2 are based on measurements by sonic rangers and pressure transducers, supplemented by stake readings. the net ablation values are calculated as the height difference between the end-of-melt-season surface level in the given year and the year before. the geographical distribution and annual net ablation records in table 2 show that for all station pairs, the lower stations measure larger ablation totals than the upper ones, because the temperature decreases with elevation and the seasonal averaged surface albedo decreases towards the ice margin. there are considerable latitudinal differences as well, as the greenland ice sheet is more than 2000 km long. ablation totals in the southern part of greenland typically amount to 3–7 m (at the lower tas, qas and nuk stations), whereas ablation totals at the more northerly sco_l and upe_l stations only amount to 2–3 m at low altitudes (<500 m a.s.l.). the ablation totals (negative value is net accumulation) from the upper stations (>500 m a.s.l.) typically amount to –0.3 to 4 m in the south (tas, qas, nuk, kan) and –0.1 to 3 m in the north (kpc, sco, upe). the record-warm year 2010 (in most regions of greenland) showed the largest ablation at most aws sites, followed by 2008, 2011, and lastly 2009 in the four-year promice record. the 2010 ablation year was characterised by relatively low winter accumulation and a long period of positive air temperatures, which are both important for net ablation (tedesco et al. 2011; van as et al. 2012). however, as the energy from solar radiation is the main contributor to the melting of snow and ice, the direct impact on ablation of inter-annual variability of temperature may be small. it was the combination of low winter accumulation and high temperatures, causing low albedo in large sections of the greenland ice sheet in 2010, that resulted in large net ablation records (tedesco et al. 2011; van as et al. 2012). the southern (qas) and south-western (nuk and kan) parts of greenland experienced a particularly anomalous ablation season, with ablation totals of up to 3 m more than in other years. even more significant was the melt response at stations close to the equilibrium-line altitude (zero net ablation) in the south and west; qas_u went from a 0.3 m surface rise in 2009 to a 3.4 m surface lowering in 2010, and at kan_m ablation increased from 0.3 m to 2.6 m during the same years. even in normal years, the kan_u station, which is placed in the accumulation zone, showed a surface lowering due to melt over the year 2010. on the other hand, the northern kpc_u station did not show anomalous ablation, as temperatures did not exceed their normal summer values (table 2). north-east greenland did not experience the anomalous circulation pattern that, in particular, southern and western greenland experienced (tedesco et al. 2011). fifteen out of 18 promice stations were successful in obtaining ablation station\year 2008 2009 2010 2011 kpc_l –* –† –† –† kpc_u –* –0.1‡ 0.1‡ 0.2‡ sco_l –* 2.6 3.5 3.1 sco_u –* 1.4 2.5 2.1 tas_l 3.6 7.4** 3.4‡ tas_u 3.3 1.5‡ 3.9$ 2.9 qas_l 7.3 – 9.3 5.5 qas_u –* –0.3‡ 3.4 0.8 nuk_l 5.3 4.8 7.2 – nuk_u 2.2 1.5 2.5 2.3 nuk_n – – –* 5.1 kan_l –* 3.5 5.4 4.0 kan_m –* 0.3‡ 2.6‡ 1.7‡ kan_u – –* – –0.2‡ upe_l – –* 3.2 2.9 upe_u – –* 2.7 2.6 thu_l – – –*† –† thu_u – – –*† –† table 2. yearly ablation totals, given as the difference between the surface height at end of the melt season and that of the previous years, in metres snow/ice equivalent per year * partial data coverage due to station establishment that year. † data may become available after the next maintenance visit. ‡ sonic ranger measurement. $ sonic ranger and stake measurements. ** sonic ranger and stake measurements for both 2009 and 2010. 7676 totals up until 2011. the remaining three stations (kpc_l, thu_l, and thu_u) did not transmit any data, but data may become available during the next maintenance visit. we obtained 42 interannual ablation totals for all the stations out of 50 possible totals. when disregarding the three stations that did not transmit, the number would be 45. this gives a success rate of 84% for all 18 stations, or 93% when disregarding the 3 non-transmitting stations. conclusions ablation totals at low altitudes (<500 m a.s.l.) typically amount to 3–7 m in the southern part of greenland, with an ablation total at the more northerly stations of 2–3 m. the ablation totals from the upper stations above 500 m a.s.l., typically amount to –0.3 to 4 m in the south (tas, qas, nuk, kan) and –0.1 to 3 m in the north (kpc, sco, upe). the interannual ablation totals show that 2010 was a year with large melt when compared to the other years. the 2010 season showed record-setting ablation totals for the southern (qas) and south-western stations (nuk and kan), however, a longer time series is needed to quantify the anomalous 2010 ablation totals (van as et al. 2012). the promice station network has successfully obtained ablation totals at the end of 2011 for 15 out 18 stations. the interannual success rate was 84% for all 18 stations and 93% if the three non-transmitting stations (kpc_l, thu_l, and thu_u) are omitted. acknowledgements the programme for monitoring of the greenland ice sheet (promice) is funded by the geological survey of denmark and greenland (geus) and the danish ministry of climate, energy and building under danish cooperation for environment in the arctic (dancea), and is conducted in collaboration with the national space institute (dtu space) and asiaq (greenland survey). the nuk and kan stations are (co-)funded by the greenland climate research centre (gcrc) and the greenland analogue project (gap), respectively. references ahlstrøm, a.p. & promice project team 2008: a new programme for monitoring the mass loss of the greenland ice sheet. geological survey of denmark and greenland bulletin 15, 61–64. bøggild, c.e., olesen, o.b., ahlstrøm, a.p. & jørgensen, p. 2004: automatic glacier ablation measurements using pressure transducers. journal of glaciology 50(169), 303–304. dahl-jensen, d. et al. 2009: the greenland ice sheet in a changing climate: snow, water, ice and permafrost in the arctic (swipa), 115 pp. oslo: arctic monitoring and assessment programme (amap). ipcc 2007: intergovernmental panel on climate change (ipcc) fourth assessment report (ar4), climate change 2007. 4 volumes. cambridge: cambridge university press. khan, s.a., wahr, j., bevis, m., velicogna, i. & kendrick, e. 2010: spread of ice mass loss into northwest greenland observed by grace and gps. geophysical research letters 37, l06501, http://dx.doi. org/10.1029/2010gl042460 rignot, e., velicogna, i., van den broeke, m.r., monaghan, a. & lenaerts, j. 2011: acceleration of the contribution of the greenland and antarctic ice sheets to sea level rise. geophysical research letters 38, l05503, http://dx.doi.org/10.1029/2011gl046583 tedesco, m., fettweis, x., van den broeke, m.r., van de wal, r.s.w., smeets, c.j.p.p., van de berg, w.j., serreze, m.c. & box, j.e. 2011: the role of albedo and accumulation in the 2010 melting record in greenland. environmental research letters 6, 014005, http://dx.doi. org/10.1088/1748-9326/6/1/014005 van as, d., fausto, r.s. and the promice project team 2011: programme for monitoring of the greenland ice sheet (promice): first temperature and ablation records. geological survey of denmark and greenland bulletin 23, 73–76. van as, d., hubbard, a.l., hasholt, b., mikkelsen, a.b., van den broeke, m.r. & fausto, r.s. 2012: large surface meltwater discharge from the kangerlussuaq sector of the greenland ice sheet during the recordwarm year 2010 explained by detailed energy balance observations. the cryosphere 6, 199–209. van den broeke, m., bamber, j., ettema, j., rignot, e., schrama, e., van de berg, w.j., van meijgaard, e., velicogna, i. & wouters, b. 2009: partitioning recent greenland mass loss. science 326, 984–986. velicogna, i. 2009: increasing rates of ice mass loss from the greenland and antarctic ice sheets revealed by grace. geophysical research letters 36, l19503, http://dx.doi.org/10.1029/2009gl040222 authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: rsf@geus.dk research article bonow & japsen 2021: geus bulletin 45 (1). 5297. https://doi.org/10.34194/geusb.v45.5297 1 of 35 peneplains and tectonics in north-east greenland after opening of the north-east atlantic johan m. bonow*1,2,3 , peter japsen3 1geovisiona ab, bro, sweden, 2department of social and economic geography, uppsala university, uppsala, sweden, 3geological survey of denmark and greenland (geus), copenhagen, denmark. abstract elevated plateaus with deeply incised valleys characterise elevated, passive continental margins (epcms) in all climate zones. these features are, however, a topic of debate regarding when and how the large-scale landscapes formed. we have investigated and mapped the partly glaciated landscape of north-east greenland (70–78°n). the area consists of crystalline basement and palaeozoic–mesozoic rift basins, capped by palaeogene basalts that erupted during the northeast atlantic break-up. our stratigraphic landscape analysis reveals a typical epcm dominated by two elevated erosion surfaces, extending 200 km east–west and 900 km north–south. the low-relief upper planation surface (ups; c. 2 km above sea level) cuts across basement and palaeogene basalts, indicating that it was graded to base level defined by the atlantic ocean in post-basalt times and subsequently uplifted. the ups formed prior to the deposition of mid-miocene lavas that rest on it, south of the study area. in the interior basement terrains, the lower planation surface (lps) forms fluvial valley benches at c. 1 km above sea level, incised below the ups. the lps is thus younger than the ups, which implies that it formed post mid-miocene. towards the coast, the valley benches merge to form a coherent surface that defines flat-topped mountains. this shows that the lps was graded to near sea level and was subsequently uplifted. hence, both the ups and the lps formed as peneplains – erosion surfaces graded to base level. the fluvial valley benches associated with the lps further indicates that full glacial conditions were only established after the uplift of the lps in the early pliocene (c. 5 ma). the uplift of the lps led to re-exposure of a mesozoic etch surface. we conclude that episodes of late neogene tectonic uplift shaped the stepped landscape and elevated topography in north-east greenland. 1 introduction elevated plateaus (or planation surfaces) that extend over wide areas are characteristic features of elevated, passive, continental margins (epcm) ( jessen 1943; king 1967; lidmar-bergström et al. 2000; japsen et al. 2012a; green et al. 2013). they occur in all climate zones from the arctic to tropical and along with mesozoic and palaeogene rift systems, along the margins of the atlantic ocean from south africa and brazil to norway and greenland. the formation of these surfaces, and when and how they reached their present elevation, is a topic of debate. are the elevated plains remnants of pre-rift *correspondence: johan.bonow@ geovisiona.com received: 20 aug 2019 accepted: 06 aug 2020 published: 21 jan 2021 keywords: cenozoic, denudation chronology, passive margin, stratigraphic landscape analysis, uplift abbreviations: afta: apatite fission-track analysis a.s.l.: above sea level aster gdem: advanced space-borne thermal emission and reflection radiometer global digital elevation model epcm: elevated passive continental margins es: etch surface globe: global land one-kilometre base elevation lps: lower planation surface odp: ocean drilling program ups: upper planation surface geus bulletin is an open access, peerreviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution and reproduction for any purpose, even commercial, provided proper citation of the original work is given. author(s) retain copyright. edited by: catherine jex (geus, denmark) reviewed by: adrian hall (stockholm university, sweden), jean-pierre peulvast (university of sorbonne, france) funding: see page 30 competing interests: none declared additional files: see page 30 https://doi.org/10.34194/geusb.v45.5297 https://orcid.org/0000-0003-0547-056 https://orcid.org/0000-0003-1689-7820 mailto:johan.bonow@geovisiona.com mailto:johan.bonow@geovisiona.com bonow & japsen 2021: geus bulletin 45 (1). 5297. https://doi.org/10.34194/geusb.v45.5297 2 of 35 www.geusbul let in.org landscapes that have remained largely unaffected by later processes (ollier & pain 1997)? did they form during rifting or break-up and remain at high elevation (gilchrist & summerfield 1990; gallagher et  al. 1998; bishop 2007; braun 2018)? did they form by glacial and periglacial processes at high elevation in arctic regions (steer et  al. 2012; egholm et  al. 2017)? or were they graded to the base level of the adjacent sea by fluvial erosion long after rifting and break-up, and were subsequently uplifted to their present elevation (lidmar-bergström et al. 2000, 2013, 2017; bonow et al. 2006a, 2006b, 2014; japsen et al. 2012a, 2019; green et al. 2013)? bonow et  al. (2014) used stratigraphic landscape analysis (lidmar-bergström et al. 2013, 2017) to identify and map two elevated planation surfaces, the upper and lower planation surface (ups and lps, respectively) in southern–east greenland (68–71°n; here referred to as the blosseville kyst region, primarily south of scoresby sund). the bedrock of blosseville kyst is dominated by flood basalts that erupted around the paleocene–eocene transition during the break-up of the north-east atlantic (larsen et al. 1989; pedersen et al. 1997; gaina et al. 2017). as these planation surfaces cut across the palaeogene basalts as well as the metamorphic basement, bonow et  al. (2014) inferred that they were graded towards the level of the newly formed north-east atlantic. consequently, the present elevation of the ups and the lps that reach 3 and 2 km fig. 1 bedrock topography of greenland: areas investigated by stratigraphic landscape analysis are indicated (bonow et al. 2006a, 2006b, 2014; and this study). the landscape of east greenland is generally of higher elevation than west greenland, where elevated areas are restricted to the coast. the load of the greenland ice sheet causes up to 850 m subsidence of the bedrock topography in central greenland. peripheral bulging along the margins of greenland, caused by this ice loading, has a negligible effect on elevation (medvedev et al. 2013). elevation data is from amante & eakins (2009). this study bonow et al. 2014bonow et al. 2006b 60°w 50°w 40°w 30°w 20°w 60°n 65°n 70°n 75°n 80°n bonow et al. 2006a elevation (m) 3000 2000 1000 –3000 –4000 –2000 –1000 0 https://doi.org/10.34194/geusb.v45.5297 http://www.geusbulletin.org bonow & japsen 2021: geus bulletin 45 (1). 5297. https://doi.org/10.34194/geusb.v45.5297 3 of 35 www.geusbul let in.org above sea level (a.s.l.), respectively, in southern east greenland, reflect uplift after their formation. three major phases of uplift and erosion led to the formation and subsequent uplift of these surfaces in late eocene, late miocene and early pliocene as estimated from apatite fission-track analysis (afta) data and landscape observations (japsen et al. 2014). these results are further confirmed by afta data from north-east greenland (japsen et al. in press). here, we continue the work of bonow et al. (2014) by presenting coherent maps and analyses of the largescale landscape of north-east greenland, north of scoresby sund (70–78°n; fig. 1). whereas the previous study focussed on the region dominated by the palaeogene basalts, this study investigates the geomorphology along the late palaeozoic – mesozoic rift system of east greenland (surlyk 1977, 1978, 2003; parsons et  al. 2017). an important question is: does the largescale landscape here mainly contain elements that are inherited from the time of rifting and break-up, or is the landscape instead dominated by younger features? for example, swift et  al. (2008) argued that the first-order topography in north-east greenland had existed since at least the time of break-up, at c. 55 ma. we apply the stratigraphic landscape analysis to identify and map planation surfaces in north-east greenland and produce a relative tectonic event chronology that defines the major phases of denudation and uplift that led to the formation of the present landscape. the results may thus provide further insight into highly debated topics in recent years, including: 1. the origin of elevated plateaus along passive continental margins (lidmar-bergström et al. 2000; japsen et al. 2009, 2012a, 2012b, 2019; hetzel et  al. 2011; green et al. 2013, 2018; haider et al. 2013; calvet et al. 2015; braun 2018; da silva et al. 2018; guillocheau et al. 2018). 2. the development of the margins of the north-east atlantic (japsen & chalmers 2000; nielsen et al. 2009; chalmers et  al. 2010; pedersen et  al. 2012; japsen et al. 2013, 2014; lidmar-bergström et al. 2013, 2017; bonow et al. 2014; egholm et al. 2017). 3. the role of tectonics for triggering the formation of the greenland ice sheet (pedersen & egholm 2013; solgaard et  al. 2013; steinberger et  al. 2015; pérez et al. 2018). 2 stratigraphic landscape analysis for mapping erosion surfaces in east greenland in this study, we apply the same technique for mapping erosion surfaces as bonow et  al. (2014), that is, stratigraphic landscape analysis (green et al. 2013; lidmar-bergström et al. 2013). we define an erosion surface graded to base level as a peneplain, in agreement with the original idea of davis (1899), who stressed the fundamental importance of a base level for the development of an eroding landscape. thus, we use the term peneplain for any erosion surface graded to base level. stratigraphic landscape analysis is based on: 1. the relationship between peneplains in crystalline basement and their cover rocks of different ages. 2. the cross-cutting relationships between such reexposed peneplains and epigene peneplains (peneplains that have never been covered by sedimentary rocks). 3. the occurrence of valleys incised below peneplains (fig. 2). stratigraphic landscape analysis is thus a further development of a long tradition in geomorphology focussed on the study of large-scale landforms that contain information about long-term erosional processes and tectonic events (davis 1899; reusch 1901; ahlmann 1919, 1941; penck 1924; baulig 1935; jessen 1943; king 1967; brunsden 1993; ahnert 1998; godard et  al. 2001; benito-calvo & pérez-gonzález 2007; peulvast et  al. 2009, 2011; hetzel fig. 2 conceptual diagram illustrating the formation of peneplains through time: a: initial topography. b: formation of a first peneplain by planation of the landscape to base level (sea level) in a tectonically stable environment. c: tectonic uplift or a significantly lowered base level results in valley incision below the peneplain. d: erosion continues within the valleys resulting in valley widening and the formation of a second peneplain controlled by the new base level. erosion primarily affects the older, elevated peneplain along its edges. e: renewed uplift ends the formation of the second peneplain and valleys again grade the landscape to the new base level. the result is a landscape with distinct steps. at this scale and for the time span considered (c. 10 myr), the downwearing of the peneplains is negligible (fu et al. 2019). d e initial topography sea level time a b c planation to base level uplift and incision of valleys valley widening incision and stepped topography https://doi.org/10.34194/geusb.v45.5297 http://www.geusbulletin.org bonow & japsen 2021: geus bulletin 45 (1). 5297. https://doi.org/10.34194/geusb.v45.5297 4 of 35 www.geusbul let in.org et al. 2011; haider et al. 2013; li et al. 2014; ma et al. 2020). classical models of landscape development focus on the idea of continuous uplift, interrupted by periods of quiescence (davis 1899; baulig 1935; fairbridge & finkl 1980; benito-calvo & pérez-gonzález 2007). however, stratigraphic landscape analysis emphasises that the relationship between relief in basement and cover rocks provides information about both uplift and subsidence of a region (lidmarbergström et al. 2013). it is important to understand the relationship between the geology of the rocks into which a surface is eroded and the present appearance and extent of the surface. an extensive erosion surface across resistant rocks (such as crystalline basement) will require a long time to form, but once formed it will persist for a long time in the landscape. on the other hand, an erosion surface that forms across less resistant rocks (such as sediments) will form more quickly, and the surface will be dissected faster than a surface formed in crystalline basement (simon-coinçon et al. 1997; fjellanger & etzelmüller 2003; bonow et al. 2009; green et al. 2013). another aspect that must be considered during the mapping of peneplains is the possible complications from structural control in terrains with flat-lying cover rocks, which are common in northeast brazil (peulvast & bétard 2015) and also in parts of north-east greenland. the formation of regional, extensive erosional surfaces is likely to be governed by the common base level to which rivers erode, given time and sufficient tectonic stability. thus, if an erosion surface loses its contact with the base level to which it formed, it is a palaeosurface, and it will begin to be dissected by incising valleys and the relief will rejuvenate. the valleys will eventually be widened, which results in a new erosion surface graded towards the base level (ahnert 1998; bonow 2004; lidmar-bergström et  al. 2017). such landscapes will thus be characterised by plateaus in distinct steps and deeply incised valleys (fig. 2), similar to the landscapes in eastern australia and southern norway, the shillong plateau, india and west greenland (lidmar-bergström et al. 2000; bonow et al. 2006a; biswas et al. 2007). the parsimonious explanation is that these plateaus are erosion surfaces that were graded to distinct base levels (japsen et al. 2009). the next parsimonious explanation is that the base levels correspond to sea level at the time of erosion of the surface. where a study area is known to have been near the sea at the time in question, and the surface is not defined by a resistant level, this is the obvious explanation. an erosion surface may be buried after its formation and thus, be preserved below its cover. such buried surfaces can be identified as unconformities in boreholes, on offshore seismic profiles and on exposed sections (e.g. larsen et al. 1989; bate 1997; lassen & thybo 2012; parsons et al. 2017). however, uplift events may lead to re-exposure of previously buried surfaces (lidmar-bergström 1989; bonow 2005; peulvast et al. 2011). the characteristics of a re-exposed surface (e.g. relative relief outliers and saprolite types) often allow it to be followed along topographical profiles or identified away from its cover. if the re-exposed surface is tilted, its extension may at some point be cut-off by younger erosion that forms a new, more horizontal surface with low-angular unconformity. in such a setting, it can be inferred that the erosion responsible for the formation of the younger surface must have involved both erosion of former cover rocks and of the re-exposed surface (fig. 3, lidmar-bergström et al. 2017). the cross-cutting relationships between peneplains of different tilts and with different relief and cover provide information about the relative denudation chronology. a relative denudation chronology is a key input to studies of uplift and erosion along a passive margin, similar to other independent datasets, such as evidence from the stratigraphic record and thermochronological data. the chronology provides tectonic information during periods where little or no geological information is otherwise available. such periods, with little or no basement mesozoic and younger sediments former base-level covered sub-mesozoic peneplain kaolinite re-exposed sub-mesozoic peneplain sea level post-mesozoic peneplain fig. 3 relationship between peneplains and cover rocks in southern sweden: the presence of mesozoic outliers and remnants of kaolinitic saprolites at a high position in the landscape, demonstrates that mesozoic cover was once more extensive. the near horizontal peneplain cut off the tilted sub-mesozoic peneplain and is therefore younger. the geological constraints and the appearance of the two surfaces reveal a history of erosion (formation of the sub-mesozoic peneplain), subsidence and deposition of the mesozoic strata, followed by tilting and uplift (change from a near-horizontal surface to an inclined surface), erosion (removal of cover and formation of a new peneplain) and a late uplift phase that explains the landscape configuration. based on lidmarbergström (1982, 1988); green et al. (2013); lidmar-bergström et al. (2013, 2017). https://doi.org/10.34194/geusb.v45.5297 http://www.geusbulletin.org bonow & japsen 2021: geus bulletin 45 (1). 5297. https://doi.org/10.34194/geusb.v45.5297 5 of 35 www.geusbul let in.org preserved geological record, often represent a larger timespan than that for which a stratigraphic record exists (e.g. ager 1973; green et al. 2013). 3 study area 3.1 geologic and tectonic setting the bedrock of north-east greenland consists of precambrian and caledonian basement overlain by palaeozoic, mesozoic and cenozoic cover rocks (fig. 4, henriksen et  al. 2008, 2009). this geological setting is favourable for an analysis of the relationship between basement and cover rocks, because the minimum age of re-exposed basement surfaces can be constrained by the age of the cover rocks, and the maximum age of surfaces that were never covered (epigene surfaces) can be constrained by the age of the cover rocks that they cut across. it is then possible to use such relationships to establish a chronology for the development of palaeosurfaces (e.g. lidmar-bergström 1988; bonow 2005; lidmar-bergström et al. 2013). the caledonian orogeny lasted from 465 to 400 ma (middle ordovician to early devonian), and culminated in the collision between laurentia and baltica about 420 ma (latest silurian; henriksen 2008). the caledonian mountains collapsed between 400 and 355 ma (early devonian to earliest carboniferous), and the sediments from the eroding caledonides were deposited in sedimentary basins parallel to the present coast in central east greenland (larsen & bengaard 1991; higgins et  al. 2008). the sea transgressed the eroded caledonian basement in the late permian (haller 1971; surlyk 1990). a series of carboniferous–mesozoic rift basins kuhn ø store koldewey dove bugt germania land hold with hope clavering ø wollaston forland d f hudson land pd m f kejser franz josephs fjord geographical society ø stauning alper nathorst land milne land renland nordvestfjord gåseland liverpool land traill ø jameson land sa f kong oscar fjord arnold escher land ymer nunatak andrée land geikie plateau scoresby sund sabine øpayer land 50 km 30°w 20°w 70 °n 72 °n 74 °n 76 °n 20°w30°w 70 °n 72 °n 76 °n 74 °n pre-devonian rock faults devonian sediment carboniferous–cretaceous sediment palaeogene basalt cenozoic intrusion cenozoic sediment quaternary sea ice n fig. 8fig. 6 fig. 4 geology of the study area (70– 78°n; location in fig. 1): df: dombjerg fault. saf: stauning alper fault. pdmf: post-devonian main fault. we refer to the pdmf and the saf as the post-devonian main fault system. crystalline basement (precambrian and caledonian) and devonian rocks dominate west of the pdmf with carboniferous– cretaceous rift basins to the east. palaeogene flood basalts dominate much of the southern part of the study area, but patches of basalt occur as far north as 75°n. several palaeogene intrusive centres occur along the coast. black dashed lines: locations of figs 6, 8. https://doi.org/10.34194/geusb.v45.5297 http://www.geusbulletin.org bonow & japsen 2021: geus bulletin 45 (1). 5297. https://doi.org/10.34194/geusb.v45.5297 6 of 35 www.geusbul let in.org developed in east greenland, forming n–s-trending, coast-parallel depocentres located to the east of the so-called post-devonian main fault (vischer 1943) – a north–north-east-trending fault, which appears as a marked scarp in the terrain, defining the present-day western boundary of the sedimentary basins north of kong oscar fjord. south of the fjord, it is termed the stauning alper fault. we will refer to these two faults as the post-devonian main fault system. important phases of rifting took place during the early and late carboniferous, late permian, late jurassic and cretaceous, prior to the opening of the north-east atlantic at the paleocene–eocene transition (surlyk 1990, 2003; stemmerik et al. 1993; stemmerik 2000; surlyk & ineson 2003). parsons et al. (2017) studied the geology of traill ø and geographical society ø and identified three main rift phases during the devonian–triassic, jurassic–cretaceous and palaeogene. the greatest amounts of faulting and block rotation occurred during palaeogene rifting, which they related to the break-up at 56 ma and to the plate reorganisation at 36 ma. break-up in the north-east atlantic was accompanied by extrusion of voluminous flood basalts that dominate the region south of scoresby sund, along blosseville kyst and its hinterland (68–70°n), where much of the lavas are referred to as the main basalts (larsen et al. 1989; brooks 2011). palaeogene basalts are also present as far as 75°n (larsen et al. 2014). the ages of the lava series near the coast in north-east greenland range from 56 to 53 ma, whereas the ages for lavas on inland nunataks in arnold escher land range from 53 to 50 ma (larsen et al. 2014). for comparison, the main basalts along blosseville kyst erupted during a time span of only one million years (56–55 ma; brooks 2011). the flood basalts along blosseville kyst attained a total vertical thickness of up to 5.5 km, but subsidence kept pace with the thickness of extruded basalts. as such, individual lava flows can be traced over thousands of square kilometres, indicating a largely horizontal lava plain without significant relief (larsen et al. 1989; pedersen et al. 1997). marine incursions onto the earliest and latest basalt flows along blosseville kyst show that the landscape was lowlying during the volcanic eruptions (wager & deer 1939; nielsen et al. 1981; larsen et al. 1989, 2013; pedersen et al. 1997; larsen & tegner 2006; brooks 2011). a phase of middle miocene volcanism is documented by the presence of the lava flows of the vindtop formation (c.  14–13 ma; storey et  al. 2004) that crop out on nunataks within a small area at 2.7–2.9 km a.s.l., just south of our study area (c. 69°n). pliocene–pleistocene deposits within our study area are reported from jameson land, île de france (c. 78°n) and store koldewey (freylinghansen et al. 1983; bennike et al. 2002, 2010). interpretation of seismic data off east greenland has provided evidence for distinct tectonic activity along the margin during the late eocene – early oligocene and mid-late miocene (larsen et  al. 1994a; hamann et  al. 2005; døssing et al. 2016; petersen 2019). 3.2 glacial history stratigraphically extensive, ice-rafted debris, including macroscopic drop stones, occur in late eocene to early oligocene sediments from the norwegian–greenland sea, indicating sediment rafting by continental ice and east greenland as the likely source (eldrett et al. 2007). eldrett et al. (2009) presented climate estimates for the eocene–oligocene based on spore and pollen assemblages in marine sediments from the norwegian–greenland sea. the climate estimates indicated cooling across the eocene–oligocene transition, but also provide evidence for relatively warm summer temperatures at that time, and thus that continental ice on east greenland was probably restricted to alpine outlet glaciers. a phase of exhumation in east greenland starting near the eocene–oligocene transition was defined from thermochronology data and was argued to have been caused by glacial erosion (bernard et  al. 2016). however, the timing is consistent with the late eocene phase of uplift and erosion defined by afta data in west, south-east and north-east greenland (japsen et al. 2006, 2014, in press). this phase of exhumation in greenland coincides with pronounced magmatic activity in east greenland (larsen et al. 2014), cessation of sea-floor spreading west of greenland and with a major plate reorganisation in the north-east atlantic (gaina et al. 2009) and is thus likely to be of tectonic origin. studying ice-rafted debris from drill cores off southeast greenland, larsen et  al. (1994a) concluded that full glacial conditions were established in south-east greenland at 7 ma. similar evidence was identified in the borehole at site 987 of the ocean drilling program leg 162 from the basin off scoresby sund. here, upper miocene – lower pliocene sediments contain abundant ice-rafted debris and evidence for slumping and turbidity currents (jansen et al. 1996; channell et al. 1999). pérez et  al. (2018) carried out seismo-stratigraphic analyses of miocene to recent deposits on the continental shelf off blosseville kyst to liverpool land. eight stratigraphic units were tied to the ocean drilling program site 987. the formation of the oldest sedimentary unit (unit 8, between the early miocene oceanic crust and a 7.3 ma discontinuity) was controlled by major tectonic events along the margin, notably the late miocene uplift episode at c. 10 ma, as defined by afta data in rock samples from the blosseville kyst region (japsen et  al. 2014). pérez et  al. (2018) explained the https://doi.org/10.34194/geusb.v45.5297 http://www.geusbulletin.org bonow & japsen 2021: geus bulletin 45 (1). 5297. https://doi.org/10.34194/geusb.v45.5297 7 of 35 www.geusbul let in.org high sediment input in this time interval by the presence of fluvial systems onshore. the subsequent, upper miocene to lower pliocene unit 7 (7.3–4.9 ma) forms a widespread sedimentary body with isolated depocentres along the outer shelf, mostly off the major inland fjords. pérez et al. (2018) inferred that cross-shelf glaciation began to influence the shelf during this interval at the time of the proposed onset of widespread glaciation in greenland around 7 ma, indicated by the occurrence of ice-rafted debris (larsen et al. 1994a, 1994b; jansen et  al. 1996; channell et  al. 1999). subsequently, two major phases of ice-sheet advance occurred across the shelf. first one was in the early pliocene (unit 6), likely influenced by topographic forcing during early pliocene uplift (japsen et al. 2014). second one was around the pliocene– pleistocene transition (unit 4). the deposits of the intervening phase (unit 5) reflect glacial retreat during the mid-pliocene warm period (raymo et  al. 1996). sedimentary successions from the pleistocene to present (units 4–1) contain features that indicate that full-scale greenland glaciation was established at 2.9 ma (sarnthein et al. 2009; pérez et al. 2018). biomolecules from the silty section at the base of deep ice cores in central southern greenland suggest that the region was forested sometime within the past million years (willerslev et al. 2007). while measurements of cosmic-ray-produced isotopes in a bedrock core from central greenland further indicate that greenland was deglaciated for extended periods during the pleistocene epoch (2.6 ma to 11.7 ka; schaefer et al. 2016; see also solgaard et al. 2013). the greenland ice sheet that today has a thickness of up to 3 km, thus, appears to have had a discontinuous history. 3.3 large-scale landscapes here, we briefly review previous work on the large-scale landscapes in east greenland. altitudes in north-east greenland vary significantly from a relatively low-lying terrain along the coast, across the carboniferous–palaeogene basins, to the elevated plains across crystalline basement, devonian sediments and palaeogene basalts further inland (fig. 4). elevation in the coastal zone rarely exceeds 1 km a.s.l.; however, the plateaus in the hinterland are typically around 2 km a.s.l., although stauning alper has peaks of up to 2.8 km a.s.l. (figs 4, 5). the plateaus are often covered by thin ice, but north of 74°n the ice sheet  almost reaches the coast. the elevation contrast between the exhumed basins along the coast and the interior basement terrains is pronounced along the post-devonian main fault system. a peneplain overlain by upper permian, shallow marine sediments, has been identified both west and east of jameson land (haller 1971; larsen 1988; surlyk 1990; vigran et al. 1999; stemmerik 2000). surlyk (1990) concluded that the peneplain represents the latest carboniferous to early permian unconformity that marks the most profound change in tectonic style and overall depositional environment in the post-caledonian development of east greenland. this marks the transition from a long period of crustal extension to a period of subsidence governed mainly by thermal relaxation of the rifted crust. a peneplain covered by middle and upper jurassic sandstones of marine origin is exposed on kuhn ø and milne land (surlyk 2003; surlyk & ineson 2003). extensive denudation that started in the early jurassic most likely led to the formation of this surface. the landscape of the coastal areas of jameson land and liverpool land is the result of tertiary plateau uplift that resulted in partial denudation of the mesozoic basins according to peulvast (1988, 1991). peulvast (1988) described a 10 km-wide westward-sloping plateau on liverpool land that he regarded as a sub-triassic planation surface. mapping showed that it disappeared to the west below the sediments on jameson land, while the surface was obliterated near the most uplifted parts on liverpool land. he regarded the westerly tilted surface on jameson land as mainly structural, although he speculated that some near-horizontal, postmesozoic peneplains might be present on liverpool land at 800–900 m a.s.l. the unconformity between palaeogene basalts and basement rocks on milne land and gåseland was mapped by larsen et al. (1989). they concluded that the basement formed a high ridge with irregular topography and large relative relief at the time of basalt eruptions, and that the basement acted as a barrier, directing various basalts flows. the saprolites encountered at the basement-basalt contact are kaolinitic (birkelund & perch-nielsen 1976). the fjord landscapes from jameson land to hudson land was studied by swift et al. (2008) who defined the ‘first-order topography’ based on a map combining elevation and slope together with elevation profiles and geology. they identified three elevational areas reflecting the main aspects of the geology, separated by escarpments along major geological boundaries. these include: 1. areas composed of mesozoic strata to the east of the post-devonian main fault system that are generally of low elevation with low to moderate relief and gently incised. 2. areas west of the post-devonian main fault where the landscape is more elevated and steeply incised by fjords. 3. areas composed of caledonian crystalline basement in the south and east that have the highest elevations and the deepest incision. https://doi.org/10.34194/geusb.v45.5297 http://www.geusbulletin.org bonow & japsen 2021: geus bulletin 45 (1). 5297. https://doi.org/10.34194/geusb.v45.5297 8 of 35 www.geusbul let in.org a strong link between elevation and lithology was noted. in particular, the caledonian basement was shown to have formed a consistent, high-elevation, low-relief landscape at c. 2000 m a.s.l., locally incised by the fjord system. the presence of palaeogene sediments beneath palaeogene basalts (jolley & whitham 2004) indicated that the present-day first-order topography had existed prior to at least 55 ma. the conclusions of swift et  al. (2008) were thus based on the configuration of the pre-basalt surface, and particularly on a contour map of that surface on milne land (larsen et  al. 1989), which they found to mirror the present topography. finally, they found that glacial modification of the landscape is strongly influenced by first-order geology and hence lithological resistance to erosion. the large-scale landscapes of north-east greenland (72–76°n) was studied by ahlmann (1941) who demonstrated that topography could provide evidence of tectonic events in the past. his main conclusion is still important: ‘… the plateau and summit areas are the remains of what has once been a more or less uniform, high plateau rising towards the west’. he interpreted the elevated plateau there as the ‘initial topography’ from which the present relief had evolved by valley incision after uplift. ahlmann argued that the landscape in north-east greenland consisted of one plateau surface, dissected by valleys, and that the landscape was arched towards the present-day coastline. he observed that the plateau cuts across the basalt sequences and therefore concluded that it was an erosional surface that had formed after the extrusion of the volcanics. as the surface was incised by deep valleys, he deduced that the timing for the uplift to its present elevation was in the late tertiary. ahlmann’s interpretation that the dominant topography is of post-basalt age is thus at odds with that of swift et al. (2008), who considered it to be pre-basalt. the regional post-basalt plateaus mapped by ahlmann continue south of the study area, into the large-scale landscape between 68°n and 71°n, studied by bonow et al. (2014). this led to the identification of two elevated, post-basalt erosion surfaces of regional extent – the ups and lps. bonow et  al. (2014) concluded that these surfaces were the result of significant erosion, a conclusion that was in agreement with studies of the zonation of zeolite minerals in the palaeogene basalts (larsen et al. 1989; neuhoff et  al. 1997). collectively, these studies showed that up to 1200 m of basalts (and possibly younger rocks) have been removed since the extrusion of the volcanics in south-east greenland. for example, at least 400 m was estimated to have been removed on milne land. 4 methods fieldwork was conducted in the study area during the summer of 2008 and 2010, using a helicopter for reconnaissance field support and as a platform for obtaining oblique photographs. we mapped planation surfaces map between 70°n and 78°n, using the methods described by bonow et  al. (2014). we used the global land one-kilometre base elevation (globe) digital elevation model for construction of a general 3d landscape model of the study area in east greenland (fig. 5) and we used the advanced spaceborne thermal emission and reflection radiometer (aster) global digital elevation model (gdem; c. 30 m resolution) for 1. construction of a map with 100 m contours (supplementary files s1, s2; fig. 6) 2. construction of topographical profiles (supplementary file s3; fig. 6) 3. construction of swath profiles (supplementary file s3; fig. 6). the elevation data contain some artefacts, for example, zero-value data occur in near-vertical mountainsides along fjords, resulting in data gaps in the model. there scoresby sund km 0 50 100 elevation (m) 2500 2250 2000 1750 1500 1250 1000 750 500 800 000 800 000 8 600 000 1 100 000 250 0 fig. 5 topography of the study area from globe 1 km data (globe task team et  al. 1999): the landscape at some distance from the coast is characterised by elevated plateaus dissected by deeply incised valleys. this configuration of landscape elements along a passive continental margin occurs on both glaciated and non-glaciated margins (e.g. lidmar-bergström et  al. 2000; green et  al. 2013). universal transverse mercator (utm) coordinates (zone 27n) shown along the map frame. https://doi.org/10.34194/geusb.v45.5297 http://www.geusbulletin.org bonow & japsen 2021: geus bulletin 45 (1). 5297. https://doi.org/10.34194/geusb.v45.5297 9 of 35 www.geusbul let in.org are also some extreme values that result in erroneous peaks in the model. further, elevation data are rather poor in areas where small, melt-water lakes appear on the ice sheet surface. topographical profiles were extracted from the aster elevation data along a square grid, spaced 25 km apart. maximum and minimum elevations along the topographical profiles were also extracted in a 50-km wide swath. the profiles were printed on paper strips to the same scale as the contour map (1:500 000). the topographical profiles in combination with the swath profiles were used to support the mapping (fig. 6). we cross-analysed 58 profiles with a cumulative length equal to 20 000 km. the mapping of the surfaces started in areas of low relative relief with only minor fluvial incision. we determined the edge of a surface where there was a rapid change of inclination as seen from closely-spaced contours on the contour map. in the low-relief areas, maximum elevation along the swath coincides with the topographical profile, and thus provides a means of expanding the surface mapping into the dissected areas. this method of combining profiles and a contour map is useful for identifying offsets and tilting within a surface (see lidmar-bergström 1988; bonow et al. 2006b, 2014). we cross-checked the interpretation made on the contour map from profile to profile to ascertain that the interpretation was consistent. we also compared the mapped surfaces with the geological maps (bengaard et al. 2007, topographical maps (geus 2007) and oblique photographs from the archives of the geological survey of denmark and greenland (geus). 40000 70000 100000 130000 scoresby sund ba 160000 190000 220000 250000 78 0 0 0 0 0 78 30 0 0 0 78 60 0 0 0 78 90 0 0 0 79 20 0 0 0 79 50 0 0 0 0 50 0 10 0 0 15 0 0 20 0 0 77 80 0 0 0 77 90 0 0 0 78 0 0 0 0 0 78 10 0 0 0 78 20 0 0 0 78 30 0 0 0 78 40 0 0 0 78 50 0 0 0 78 60 0 0 0 78 70 0 0 0 78 80 0 0 0 78 90 0 0 0 79 0 0 0 0 0 79 10 0 0 0 79 20 0 0 0 79 30 0 0 0 79 40 0 0 0 79 50 0 0 0 79 60 0 0 0050 0 10 0 0 15 0 0 20 0 0 pr ofi le n s7 0 0 topography along profile minimum topography in swath maximum topography in swath fig. 6 construction of topography profiles with minimum and maximum elevations within a swath: a: north–south profile with topography along the profile transect and with maximum and minimum elevations within the swath. x-axis: utm northing (km). y-axis: elevation (km). location in fig. 4. b: the 100-m contour map constructed from aster data. these data are used in the surface mapping and to construct topography profiles. utm coordinates indicated (km; utm zone 27n). grey area: 50-km wide swath used to calculate maximum and minimum elevation along the profile. black dashed line: the actual topography along profile shown in panel a. https://doi.org/10.34194/geusb.v45.5297 http://www.geusbulletin.org bonow & japsen 2021: geus bulletin 45 (1). 5297. https://doi.org/10.34194/geusb.v45.5297 10 of 35 www.geusbul let in.org 5 results this study overlaps with the mapping of bonow et  al. (2014) in the area between 70°n and 71°n, where we identify the same elevated landscape features – the ups and lps (figs 7, 8). we see that these features extend across northeast greenland between 70°n and 78°n (figs 5, 9), and so they are the focus of the mapping and analysis in the current study. we also report the mapping of an exhumed etch surface (es) formed by deep weathering (supplementary files s1, s2). high-resolution contour maps for the northern and southern parts of the study area, overlain with the outline of the erosion surfaces, are provided as supplementary files s1 and s2, respectively. topographic profiles are also provided as a supplementary file s3. 5.1 mapping of erosion surfaces 5.1.1 the deeply weathered basement around milne land and gåseland, weathered basement rocks are exposed at their contact with both palaeogene basalts (fig. 10) and jurassic sediments (fig. 11). a detailed analysis of these sub-palaeogene and sub jurassic, weathered surfaces is beyond the scope of this paper. however, these surfaces are important as they provide snapshots of the erosional and depositional pre-basalt history of the margin. for example, the presence of weathered basement that crops out below marine jurassic sediments shows that after the period of weathering and erosion of the land surface, prior to the jurassic, the landscape was buried during subsidence, resulting in deposition of the marine sediments. a later ? ? ? ? ? ? ? ? ? ? ups, mapped and intepreted store koldewey clavering ø d f pd m f stauning alper milne land gåseland sa f kong oscar fjord ymer nunatak pre-devonian rock faults devonian sediment carboniferous–cretaceous sediment palaeogene basalt cenozoic intrusion cenozoic sediment quaternary sea ice renland a scoresby sund n 50 km 30°w 20°w 70 °n 72 °n 74 °n 76 °n 20°w30°w 70 °n 72 °n 76 °n 74 °n fig. 7 outlines of the upper planation surface (ups) and lower planation surface (lps) in north-east greenland: a: ups . b: lps. see fig. 4 for additional place names and supplementary files s1 and s2 for maps of these surfaces at 1:500 000 scale. https://doi.org/10.34194/geusb.v45.5297 http://www.geusbulletin.org bonow & japsen 2021: geus bulletin 45 (1). 5297. https://doi.org/10.34194/geusb.v45.5297 11 of 35 www.geusbul let in.org phase of uplift that must have occurred after the jurassic raised the marine sediments to their present position in the landscape, which are now visible overlying the weathered basement. this is geological evidence of an episodic development of the landscape, in contrast to the frequent assumption of a progressive emergence of the bedrock (e.g. pedersen et al. 2012). the outline of the weathered basement is shown in supplementary file s1, and some of the details of the weathered basement are discussed by larsen et  al. (1989). where a bedrock surface is formed by deep weathering, its character does not depend on the age of the cover rocks, for example, jurassic sediments or paleocene basalts. for the purpose of this paper, which is to distinguish between the pre-basalt and the postbasalt landscape development, we chose to refer to these different pre-basalt surfaces as a single es, as this term describes the formation process, the characteristics of saprolites and the general hilly relief shape of the basement (e.g. bonow 2005). the es can be mapped with high confidence close to cover rocks, but it is more difficult to identify in areas with no cover and further away from cover rocks. it is especially difficult to identify deeply weathered surfaces in formerly glaciated areas because classical roche moutonnée landscapes (e.g. sugden 1974; glasser & warren 1990; freire et al. 2015) are highly similar in shape to bedrock forms shaped by weathering (lindström 1988; sugden et  al. 1992; andré 2001, 2002; migoń & lidmar-bergström 2001; bonow 2005; krabbendam & store koldewey clavering ø d f pd m f geographical society ø stauning alper renland traill ø sa f kong oscar fjord pre-devonian rock faults devonian sediment carboniferous–cretaceous sediment palaeogene basalt cenozoic intrusion cenozoic sediment quaternary ice sea scoresby sund lps, intepreted lps, mapped b n 50 km 30°w 20°w 70 °n 72 °n 74 °n 76 °n 20°w30°w 70 °n 72 °n 76 °n 74 °n fig. 7 (continues) outlines of the upper planation surface (ups) and lower planation surface (lps) in north-east greenland: a: ups . b: lps. see fig. 4 for additional place names and supplementary files s1 and s2 for maps of these surfaces at 1:500 000 scale. https://doi.org/10.34194/geusb.v45.5297 http://www.geusbulletin.org bonow & japsen 2021: geus bulletin 45 (1). 5297. https://doi.org/10.34194/geusb.v45.5297 12 of 35 www.geusbul let in.org milne land renland scoresby sund ups ups ups eses upsups lpslps gåseland 0 elevatio n (m ) elevatio n (m ) 500 1000 1500 2000 0 500 1000 1500 2000 fig. 8 3d model from aster data showing the characteristic landforms in the region around gåseland, milne land and renland: a: 3d model with no labels. b: the same 3d model with labels. the upper planation surface (ups) is well preserved, while the lower planation surface (lps) is mainly identified along the main valleys. the inclined jurassic etch surface (es) is identified in eastern milne land where it is cut off by the ups. location in fig. 4. bradwell 2014; lidmar-bergström et al. 2017; hall et al. 2020). all parts of the identified es in the study area are close to cover rocks and thus, we have high confidence in our mapping of its extent. on milne land, the tilted es is cut off by the near-horizontal ups that formed across both basement and basalts, which shows that the ups formed after the es (figs 11, 12). 5.1.2 the upper planation surface the ups dominates the landscape away from the coast, especially south of 74°n. the ups is up to 200 km wide and extends across the study area for more than 900 km in north–south direction. photographs provide further visual evidence of the continuity of the ups across wide areas (figs 11–14). north of 74°n, the ups disappears below the greenland ice sheet. the ups is to a large extent covered by ice due to its high elevation at around 2 km a.s.l. the ice increases in thickness towards the west. in the northernmost part of the study area, the ups occurs at slightly lower elevation, c. 1.5 km. south of scoresby sund the ups becomes more elevated (above 3 km) in the kangerlussuaq area (c. 400 km south of the study area; bonow et al. 2014). in the south, a flat bedrock surface was observed in section and photographed along the rim of incised valleys, emerging from beneath the thin ice caps. these plateau rims are typically too small to be represented on the maps, but these observations support the mapping and allowed us to interpret the bedrock surface as part of the ups. north of 74°n, the ups is represented by minor, flattish remnants located where the greenland ice sheet becomes coherent (fig. 7a); for example, ymer nunatak in the northernmost part of the study area (fig. 4). it is not possible to assess the relief of these surfaces as only the rim is exposed from below the ice cover, but the rims appear to be flat. the correlation between the well-developed ups in the south and the remnants in the north is less well constrained because of the ice cover and the long distance between them. thus, in this part of the study area, it is not always clear how to correlate the surface on the topographical profiles. in the south, especially on milne land and gåseland, the ups is developed across rocks of different age and of different resistance (fig. 7a). for example, on milne land the ups developed across palaeogene basalt (fig.  15a) and crystalline basement rocks (fig. 15b) just a few kilometres apart and at about the same elevation. on gåseland, the ups is about 700 m below the original top of the volcanic pile as estimated from zeolite stratigraphy (larsen et al. 1989; fig. 12), offering further evidence that the position of the ups is not defined by structural control on a regional scale from flat-lying lava flows. the ups cuts across the palaeogene basalts at about 2.2 km a.s.l. in arnold escher land, located about 200 km from the coast and close to the ice sheet (74°n, 28°w; https://doi.org/10.34194/geusb.v45.5297 http://www.geusbulletin.org bonow & japsen 2021: geus bulletin 45 (1). 5297. https://doi.org/10.34194/geusb.v45.5297 13 of 35 www.geusbul let in.org fig. 4). the ups is thus mainly preserved in basalts and basement rocks in the southern part of the study area. however, the landscapes across the sedimentary basins near the coast are characterised by many ridges and minor flattish summits that reach the projected level for the ups, even in areas such as traill ø with sedimentary rocks at 1800 m a.s.l. offsets of the ups along the topographical profiles would indicate tectonic movements after the formation of the ups. we could not identify any, but some movements cannot be ruled out as the ice covering the ups obscures its precise identification in many places, and correlating the ups across a fjord or a valley is occasionally problematic. where there are offsets, however, they are only minor, indicating that the entire region where the ups is preserved has been uplifted and appears to have acted as a single tectonic unit. a deviation from this pattern occurs in stauning alper where the summits reach an elevation of 2.8 km a.s.l. – the highest elevation within the study area. even though no remnants of the ups were identified in this area, it was possible to use the swath profiles to reconstruct the elevation of the ups fairly well (fig. 16). the contours of the ups define a dome around stauning alper, whose highest area is above c. 2.3 km a.s.l. (fig. 9a). the domal pattern of the ups in this region indicates that a fairly localised, tectonic movement occurred after the formation of the ups. south of the study area, in the blosseville kyst region, the elevation of the ups increases and approaches 3.5 km a.s.l. near gunnbjørn fjeld, which at 3.7 km a.s.l. is the highest mountain in greenland (bonow et al. 2014). the ups is well defined across large areas, but less so where only a limited part of the ups is preserved, including coastal areas with only a narrow ice-free landmass. here, the precise extent and the correlation of the surface between profiles (especially in north–south direction) are uncertain. this is the case for the ups across most of the sedimentary areas in the east and in the region north of 74°n. another source of uncertainty is the presence of extensive ice caps that cover much of the surface. this problem cannot be solved completely, but the presence of minor remnants of flat surfaces or nunataks, supports the interpretation that the ups is present below the ice. this is further supported by the occurrence of outliers of palaeogene basalts on the inland nunataks in arnold escher land that are equivalent to, but slightly younger than, the basalts on milne land (fig. 4; larsen et al. 2014). in summary, the formation of the ups must have been governed by the general base level at the time of formation, as the surface cuts across rocks of different age and resistance (figs 11–13, 15). therefore, the ups is not a structural surface, such as the top of the last basalt flow, which an investigation of a minor area might suggest. since the ups developed after the extrusions of the palaeogene basalts and the onset of sea-floor spreading in the north-east atlantic, the general base level must have been the atlantic ocean. hence, we can consider the ups as a peneplain. 5.1.3 the lower planation surface the lps typically occurs at c. 1 km a.s.l. and it extends across a 50 to 75 km wide zone (up to 100 km) inland from the coast (fig. 7b). it also exists as wide-valley benches along some of the major fjords further inland (figs 17, 18). in the areas where sedimentary rocks crop out, the lps is defined by flat-topped summits that make up a coherent surface, while in areas with basement rocks, the lps is more extensive and less dissected (fig. 18). north of 74°n, the lps dominates the ice-free areas of the upper-plateau landscape in near-coastal areas (fig. 7b). however, south of 74°n, the lps and the ups co-exist, and the lps can occasionally be followed as a wide-valley bench along some of the major valleys (figs 9, 17). the lps reaches an elevation of 2 km a.s.l. in the southern part of the blosseville kyst area (c. 68⁰n; bonow et  al. 2014), where it is continuous across different rock types with significantly different resistance to erosion. the difference in geology defines how developed the lps is; that is, less developed in resistant rock and vice versa. the lps is incised below the ups, and therefore younger. the lps thus extends from valley benches in basement rocks in the interior to plateau remnants in sedimentary rock at the same elevation closer to the coast. this supports the finding that the formation of the lps was controlled by the fluvial system and thus graded towards the new base level, which again must have been the atlantic ocean, and hence we can also consider the lps to be a peneplain. 5.1.4 escarpments and faults the post-devonian main fault system is a prominent feature that runs north–south through the study area (fig. 4). east of the fault, where sediments are exposed, the ups is not present, possibly because it has been destroyed. the lps, however, is defined on both sides of the post-devonian main fault at about the same elevation (fig. 7b). west of the fault, the lps is defined by wide-valley benches, and the surface continues at the same elevation east of the fault, where it is defined by flat-topped summits. the escarpment between the ups and the lps follows along the major valleys. 5.2 relative denudation chronology and magnitude of uplift 5.2.1 constraints on the timing for the formation of the ups and lps the ups post-dates the extrusion of the basalts in the paleocene–eocene transition (c. 56 ma) as the surface truncates both the palaeogene basalts and older https://doi.org/10.34194/geusb.v45.5297 http://www.geusbulletin.org bonow & japsen 2021: geus bulletin 45 (1). 5297. https://doi.org/10.34194/geusb.v45.5297 14 of 35 www.geusbul let in.org rocks. the wide areal extent and low relative relief of the surface indicate long-term, tectonically stable conditions during its formation as, given enough time, resistant rocks will be denuded until all significant topographical expression has been removed (fig. 2). this could take several million years, as was suggested for west greenland (bonow et  al. 2006b, 2007a, 2007b; japsen et al. 2009). we therefore find that the ups represents a single erosional phase that affected east greenland, at least between 68°n and 78°n as documented in this paper and by bonow et al. (2014). the presence of the mid-miocene lava flows of the vindtop formation (14–13 ma; storey et al. 2004), south of our study area (69.18°n), provide further constraints on the timing of the formation of the ups. the vindtop formation rests on the palaeogene basalts in a small area between 2.7 and 2.9 km a.s.l. according to the mapping of bonow et al. (2014), this unconformity corresponds to the ups and this implies that the ups had formed prior to the eruption of the mid-miocene volcanics, well before the onset of the plio–pleistocene glaciations (c. 2.7 ma; shackleton et al. 1988). as the valley benches associated with the ups dominates a lps dominates clavering ø liverpool land renland store koldewey ups and lps overlap lps elevation ups elevation scoresby sund 74°n 50 km 72°n 70°n 30°w 25°w 20°w 1.5 1.0 1.1 1.1 1.1 1.0 1.2 1.1 1.1 1.3? 1.0 1.4 1.2 1.5 1.8 1.8 1.7 2.3 2.0 stauning alper fig. 9 extent and elevation of the upper and lower planation surface (ups and lps): a: ups and lps shown with elevation contours. b: ups and lps annotated with the location and orientation of figs 10–20. the ups is typically 2 km a.s.l., identified mainly south of 74°n at some distance from the coast. north of 74°n, the ups is identified only as small remnants close to the greenland ice sheet. the lps is typically 1 km a.s.l. and coalesces to form a coherent surface towards the coast. note the dome structure of the ups centred on stauning alper, which suggests tectonic doming after formation of the ups. https://doi.org/10.34194/geusb.v45.5297 http://www.geusbulletin.org bonow & japsen 2021: geus bulletin 45 (1). 5297. https://doi.org/10.34194/geusb.v45.5297 15 of 35 www.geusbul let in.org es fig. 10 the contact between palaeogene basalt and crystalline basement is the re-exposed etch surface (es), milne land: the basement is weathered and kaolinitic saprolites are found at the contact (e.g. larsen et al. 1989). photo location in fig. 9. fig. 9 (continues) extent and elevation of the upper and lower planation surface (ups and lps): a: ups and lps shown with elevation contours. b: ups and lps annotated with the location and orientation of figs 10–20. the ups is typically 2 km a.s.l., identified mainly south of 74°n at some distance from the coast. north of 74°n, the ups is identified only as small remnants close to the greenland ice sheet. the lps is typically 1 km a.s.l. and coalesces to form a coherent surface towards the coast. note the dome structure of the ups centred on stauning alper, which suggests tectonic doming after formation of the ups. 70°n 30°w 25°w 20°w 72°n 74°n 76°n ups dominates lps dominates clavering ø renland store koldewey ups and lps overlap stauning alper b 50 km 13c 13d 14a 14b 14c 11a 15 a b 12 16a 16b 16c 18a 18b 10 14d 11b liverpool land 13b 13a 20 12c kejser franz josephs fjord nathorst land furesø milne land th. sørensen land hinks land nordvestfjord tyrolerfjord flyverfjord øfjord gåsefjord gåseland fønfjord andrée land geikie plateau scoresby sund payer land dr. margrethe ii land https://doi.org/10.34194/geusb.v45.5297 http://www.geusbulletin.org ups ups scoresby sund etch surface in basement jurassic sediments jurassic sediments basement volcanics volcanics basement basement ups palaeogene volcanics jurassic sediments etch surface in basement 0 1 2 3esewnw ups es milne landrødefjordrenland jurassicups lps elevation (km) 0 100 200 300 km topography along profile minimum topography in swath maximum topography in swath a b c fig. 11 the relationship between the deeply weathered basement (etch surface, es) overlain by the middle jurassic charcot bugt formation and palaeogene basalts on southern milne land. the upper planation surface (ups) truncates the sub-jurassic es and the basalts. location of photographs and profile are shown in fig. 9. a: view from southern milne land across scoresby sund towards the geikie plateau where the ups cuts across the basalts. the tilted es of sub-jurassic age is characterised by distinct hills defined by intensively weathered fracture systems prior to the deposition of the cover rocks. jurassic sediments rest on the es, which therefore constrains the final formation age of this surface. b: view of a tilted es on southern milne land looking northwest. the extensive ups is developed across renland and beyond. the ups truncates both the basement and the volcanics, which constrains the ups as younger than the palaeogene basalts (see also fig. 15). note that jurassic sediments cover the es, but also that valleys within the es have been filled by basalt flows. this means that the jurassic es surface had been re-exposed prior to eruption of the palaeogene basalts. black line: approximate location between the 210 and 340 km marks of the profile in c. photo: kort & matrikelstyrelsen, denmark. c: swath profile illustrates the relationship between three surfaces: es, ups and the lower planation surface (lps). bonow & japsen 2021: geus bulletin 45 (1). 5297. https://doi.org/10.34194/geusb.v45.5297 17 of 35 www.geusbul let in.org b a fønfjord gåseland milne land gåsefjord geikie plateau 10 km basement milne land fm geikie plateau fm rømer fjord fm skrænterne fm cover vvvv v v v v v v nwse top of lava pile predicted from elevation of zeolite zones post-basalt offset ice surface c basement basement basalt es es ups ups ups 0 1 2 km fig. 12 the upper planation surface (ups) at an elevation of about 1800 m a.sl. developed across palaeogene basalts on gåseland. the ups was likely formed by erosion after the extrusion of the volcanics. locations shown in fig. 9. a: view to the north-west across the inner part of gåsefjord b: the volcanic strata are tilted south-east and about 700 m of basalt have been removed from above gåseland (larsen et al. 1989). the etch surface (es) defines the base of the basalts. photo: kort & matrikelstyrelsen, denmark. c: topography profile showing the inclined and truncated strata of the main basalts and the original top of the lava pile estimated from the elevation of zeolite zones across milne land, gåseland and geikie plateau. the present-day lava surface is consequently an erosional feature and the tilt and offset of the basalt formations are post-basalt features. the ups on gåseland shown in panel b is thus about 700 m below the original top of the volcanic pile. note the tilt of four formations of the main basalts, geikie plateau fm, skrænterne fm, milne land fm and rømer fjord fm. based on larsen et al. (1989) and bonow et al. (2014). https://doi.org/10.34194/geusb.v45.5297 http://www.geusbulletin.org http://a.sl bonow & japsen 2021: geus bulletin 45 (1). 5297. https://doi.org/10.34194/geusb.v45.5297 18 of 35 www.geusbul let in.org a b th. sørensen land renland renland nathorst land nordvestfjord fig. 13 the upper planation surface (ups) developed across basement rocks in the south-western interior of north-east greenland: a: looking south-west of th. sørensen land and renland across nordvestfjord. b: looking north-east from renland towards nordvestfjord and nathorst land. c: looking south-west across the southern part of stauning alper towards nordvestfjord and renland. d: looking south of nathorst land towards furesø. photo locations in fig 9. photo: kort & matrikelstyrelsen, denmark. lps are fluvial in origin, this also suggests that the lps was completed prior to the onset of the major, late cenozoic glaciations at c. 2.7 ma. the incision of the lps below the ups and the wide areal extent of the lps in the north suggest: (1) a change of base-level (uplift) after the formation of the ups and (2) re-establishment of relatively stable tectonic conditions so that the planation surface could develop across a large area. this cycle of events was repeated as the lps is also uplifted. the lps must have formed during the late neogene as it is younger than the ups that had formed by mid-miocene times. 5.2.2 amount of uplift we have shown that the ups and the lps were formed as erosion surfaces graded to former base level and therefore, their present elevation and the difference in elevation between them can be used to estimate the amount of uplift since their formation (fig. 19). the vertical distance between these two surfaces represents the magnitude of the first post-ups uplift event, allowing for subsequent erosion. the distance from the lps to the present sea level represents the second post-ups uplift event. this means that after the ups was graded to base level, an uplift phase raised it to about 1 km https://doi.org/10.34194/geusb.v45.5297 http://www.geusbulletin.org bonow & japsen 2021: geus bulletin 45 (1). 5297. https://doi.org/10.34194/geusb.v45.5297 19 of 35 www.geusbul let in.org c d nordvestfjord renland furesø nathorst land a.s.l., leading to incision that graded the lps to sea level. a subsequent uplift phase of c. 1 km raised the whole landscape, leaving the lps at about 1 km a.s.l. however, in the northernmost parts of the study area the lps did not reach such heights, for example, 600 m on germania land. these two uplift events define the main appearance of the present-day topography in northeast greenland. 5.2.3 relative denudation chronology based on the mapping of the planation surfaces and the identification of re-exposed surfaces combined with the geological record, we propose the following relative denudation chronology for north-east greenland. denudation and weathering of the basement in pre-middle jurassic and pre-eocene times led to the formation of deeply weathered basement surfaces characterised by a hilly relief. the extrusion of palaeogene basalts during the break-up of the north-east atlantic was accompanied by rapid subsidence (e.g. brooks 1985, 2011; larsen et al. 1989). the first uplift phase in post-basalt times was accompanied by erosion to the new base level, resulting in the formation of the ups as a regional denudation surface that cuts across the palaeogene basalts and older rocks. this phase most likely happened around the eocene– oligocene transition when uplift of the inner margin of south-east greenland triggered a sudden and strong fig. 13 (continues) the upper planation surface (ups) developed across basement rocks in the south-western interior of north-east greenland: a: looking south-west of th. sørensen land and renland across nordvestfjord. b: looking north-east from renland towards nordvestfjord and nathorst land. c: looking south-west across the southern part of stauning alper towards nordvestfjord and renland. d: looking south of nathorst land towards furesø. photo locations in fig 9. photo: kort & matrikelstyrelsen, denmark. https://doi.org/10.34194/geusb.v45.5297 http://www.geusbulletin.org bonow & japsen 2021: geus bulletin 45 (1). 5297. https://doi.org/10.34194/geusb.v45.5297 20 of 35 www.geusbul let in.org a b hinks land th. sørensen land nathorst land nathorst land hinks land nordvestfjord flyverfjord fig. 14 thin ice covers the upper planation surface (ups) in many locations in north-east greenland: a: view along nordvestfjord towards the south-east. b: the east–west-trending flyverfjord. north is to the left of the image. c: looking west along the inner part of kejser franz joseph fjord towards andrée land. d: view from milne land towards the south across scoresby sund where the ups defines the top of the basalts on geikie plateau. photo locations in fig. 9. photos a–c: geus archive. https://doi.org/10.34194/geusb.v45.5297 http://www.geusbulletin.org bonow & japsen 2021: geus bulletin 45 (1). 5297. https://doi.org/10.34194/geusb.v45.5297 21 of 35 www.geusbul let in.org fig. 14 (continues) thin ice covers the upper planation surface (ups) in many locations in north-east greenland: a: view along nordvestfjord towards the south-east. b: the east–west-trending flyverfjord. north is to the left of the image. c: looking west along the inner part of kejser franz joseph fjord towards andrée land. d: view from milne land towards the south across scoresby sund where the ups defines the top of the basalts on geikie plateau. photo locations in fig. 9. photos a–c: geus archive. andrée land kejser franz joseph fjord c d scoresby sund geikie plateau ättestupan https://doi.org/10.34194/geusb.v45.5297 http://www.geusbulletin.org bonow & japsen 2021: geus bulletin 45 (1). 5297. https://doi.org/10.34194/geusb.v45.5297 22 of 35 www.geusbul let in.org ups ups basalt basement a b es øfjord basement fig. 15 the upper planation surface (ups) developed at an elevation of about 1800 a.s.l. across palaeogene basalt and basement on milne land, constraining the age of the ups to be younger than the palaeogene basalts. a: ups cutting across the basalt. es: etch surface. b: ups cutting across basement, note øfjord in the background. there are only a few kilometres between the two locations. photo locations in fig. 9. 0 1000 2000 3000 topography along profile minimum topography in swath maximum topography in swath a b w ups e 0 100 200 km c fig. 16 topography across stauning alper: a, b: landscape with alpine relief and no remnants of the upper planation surface (ups) across stauning alper. note that the ups is present towards the north-west in the background of panel a. photo a: geus archive. photo b: mette olivarius. c: w–e profile illustrating that the peaks of the alpine relief can be projected to a surface (blue dashed line) that is coherent in the west. x-axis: utm easting (km). y-axis: elevation (m). locations in fig. 9. influx of coarse, clastic turbidites offshore. these turbidites overlay fine-grained lower eocene sediments (larsen et al. 1994b; bonow et al. 2014). when this planation surface had finally formed, it had a low topographical expression in the landscape, graded to near the base level at that time, which must have been the adjacent atlantic ocean. the presence of the mid-miocene vindtop formation above the palaeogene basalts just south of our study area, implies that the ups had formed by mid-miocene times and thus well before the onset of glaciation (bonow et al. 2014). glacial cirques and valleys cutting into the ups also support this observation. the development of the ups was terminated by a second, post-basalt uplift phase that raised the landscape by about 1 km. this led to incision of valleys below the uplifted ups and to the development of a new surface, the lps, during the late neogene as it is younger than the ups that had formed by mid-miocene times. the development of the lps was terminated by a third base level change in post-basalt time, due to uplift of the landscape by about 1 km, which then led to dissection of the lps. the further destruction of both the ups and lps was partly glacial and occurred during the late cenozoic glaciations. 6 discussion 6.1 does present-day topography reflect the pre-break-up landscape? new mapping of traill ø and geographical society ø indicates three main rift phases during the devonian–triassic, https://doi.org/10.34194/geusb.v45.5297 http://www.geusbulletin.org bonow & japsen 2021: geus bulletin 45 (1). 5297. https://doi.org/10.34194/geusb.v45.5297 23 of 35 www.geusbul let in.org fig. 17 relationship between the upper and lower planation surface (ups and the lps, respectively) along the north side of nordvestfjord: lps is well developed on the north side of nordvestfjord and there is a distinct escarpment towards the ups that is well developed across nathorst land. note the consistent level of the lps and that minor valleys often stop incising at that valley bench (black arrows). the inset map shows the photo location and the detailed mapping of the ups (blue) and the lps (orange) in the area, see supplementary file s1. photo: kort & matrikelstyrelsen, denmark. v ups ups nordvestfjord nathorst land lps lps lps lps jurassic–cretaceous and palaeogene with the greatest amounts of faulting and block rotation occurring during the latter phase related to break-up of the north atlantic c. 56 ma and to the plate reorganisation c. 36 ma (parsons et  al. 2017). the cross sections published by parsons et al. (2017; fig. 10) show no correlation between the present-day landscape and the underlying mesozoic sediments and the fault systems that offset them. this lack of correlation agrees well with our identification of the lps in the summits over large parts of traill ø and geographical society ø (fig. 7b). as argued in section 5.2.1, the lps was likely graded to the base level in the late neogene, and it is therefore significantly younger than the dominant palaeogene phase of block rotation. consequently, faulting prior to, during, or after break-up has left few traces in the present-day landscape. swift et al. (2008) found the first-order topography in north-east greenland to be closely related to the main aspects of the regional geology, in particular, controlled by escarpments along major geological boundaries. they also used the configuration of the pre-basalt surface from a small area in milne land in relation to the post-devonian main fault system to argue that the present first-order topography had existed since at least the time of break-up, c. 55 ma. swift et al. (2008) did not, however, consider ahlmann’s (1941) conclusion that the first-order topography must have developed in post-basalt time, following the observation that planation had occurred across both palaeogene basalts and basement rocks to form the continuous erosion surface on a regional scale, which is now present as an elevated plateau (the ups). the detailed mapping presented here also shows that there is no change in the separation between the ups and lps across the post-devonian main fault system, so the fault does not control the development of the first-order topography. instead, the step (escarpment) between the two surfaces winds along the main valleys that cut across the fault system (fig. 7b). geological evidence from the blosseville kyst region (nielsen et al. 1981) further shows that the flood basalts were extruded near sea level and the kilometre-scale subsidence kept pace with the extrusion of the basalts (see bonow et al. 2014 and references therein). the presence of late paleocene to earliest eocene marine strata at elevations up to 700 m a.s.l. on wollaston forland, hold with hope and sabine ø (nøhr-hansen et  al. 2011), demonstrates that the present-day landscape and topography is very different from the landscape at the time of continental break-up. furthermore, these observations document that post break-up rock-uplift was at least 700 m, which is in agreement with the elevation of the lps in this region of about 1 km. however, possible post-rift burial of the marine sediments would lead to a bigger magnitude of their vertical movement since maximum burial. pedersen et al. (2012) presented a different view on the origin of the present-day mountains in north-east greenland. they suggested that the mountains represent remnants of the original caledonian topography https://doi.org/10.34194/geusb.v45.5297 http://www.geusbulletin.org bonow & japsen 2021: geus bulletin 45 (1). 5297. https://doi.org/10.34194/geusb.v45.5297 24 of 35 www.geusbul let in.org a b clavering ø tyrolerfjord lps lps lps store koldewey lps payer land fig. 18 the lower planation surface (lps) in the northern part of the study area: a: view of the lps at c. 1300 m a.s.l. on payer land seen from clavering ø towards the north-west across tyrolerfjord. minor remnants of flattish land in the foreground are interpreted as the lps developed in sedimentary cover rocks. the extensive surface in the background formed across crystalline basement. photo: geus archive. b: view east, overlooking basement terrain in dronning margrethe ii land with store koldewey in the background. the flat summits are interpreted to be part of the lps and are c. 1 km a.s.l. photo: kort & matrikelstyrelsen, denmark. photo locations in fig. 9. modified during early rifting and that the area has undergone slow, steady exhumation since c. 250 ma. they based this hypothesis on inverse modelling calibrated by afta data from samples of exposed caledonian basement. however, afta data on their own register only episodes of cooling and must be integrated with geological evidence to reveal episodes of reheating that reflect re-burial. the modelling results of pedersen et  al. (2012) are at odds with the observations presented here that show two elevated plateaus, formed and uplifted after break-up, dominate the landscape of north-east greenland. japsen et  al. (2013) described how the well-documented geological record of north-east greenland (e.g. surlyk 1990) shows that the caledonian mountains were obliterated as topographic features during the late palaeozoic. the presence of outliers of both upper carboniferous and middle jurassic sediments resting on the basement in the study area of pedersen et al. (2012) also provides evidence of a history involving episodic, post-caledonian exhumation and re-burial. japsen et al. (2013) concluded that the geological record in east greenland is https://doi.org/10.34194/geusb.v45.5297 http://www.geusbulletin.org bonow & japsen 2021: geus bulletin 45 (1). 5297. https://doi.org/10.34194/geusb.v45.5297 25 of 35 www.geusbul let in.org fig. 19 conceptual model that demonstrates how to estimate the amount of uplift based on the elevation of peneplains and their relationship to the base level. ups: upper planation surface. lps: lower planation surface. amount of first post-ups uplift amount of second post-ups uplift ups lps incised valley former sea level present sea level former sea level incompatible with the idea of slow, steady exhumation since the caledonian orogeny. 6.2 formation and preservation of surfaces the characteristic basement hills, separated by minor valleys, which frequently appear below the palaeogene basalts, are the result of erosion during different weathering episodes in the mesozoic (larsen et al. 1989). at that time, the basement was exposed on the earth’s surface and became deeply weathered, forming saprolites, under warm, humid conditions. stripping of saprolites occurred during drier conditions. after its formation and until recently, this weathered basement surface (es) and the saprolites must have been protected from erosion by a cover of sediments or basalts, as seen in the area around milne land. whereas the basement areas at high elevation are planated and typically form part of the ups, the characteristic hilly relief of the es occurs only at low elevation. the es must have been covered when the planation occurred, otherwise this relief would have been obliterated by the planation (figs 3, 11), which suggests a recent episode of exhumation due to late uplift. a similar history with mesozoic deep weathering, followed by burial and late cenozoic exhumation, has also been documented in west greenland (bonow 2005), southern greenland (peulvast et al. 2011) and scandinavia (lidmar-bergström 1988, 1989; lidmar-bergström et al. 2017). the ups is preserved mainly in areas with resistant rocks such as crystalline basement and palaeogene basalts (e.g. on gåseland and milne land; fig. 12). there are only minor remnants of the ups in coastal areas underlain by sedimentary sequences, for example on traill ø. prior to the first uplift of the ups, the surface most likely extended across a much wider area towards the coast, across the sedimentary basins on jameson land, for example. independent lines of evidence support the suggestion that the ups once covered a much wider area towards the coast. palaeogene volcanics are present offshore between 72°n and 76°n (henriksen et  al. 2009), and it is thus likely that a thick cover of basalts was present in the coastal zone of the study area, where the basalts occur as outliers today (fig. 4), and across jameson land (mathiesen et al. 2000). by inference, it is likely that the ups developed across this now-lost, basalt cover. similarly, a late palaeogene unconformity mapped on seismic sections across the jan mayen micro-continent (around 68°n, 12°w), has been suggested as an equivalent of the ups onshore in east greenland (blischke et al. 2019). while the ups is preserved mainly in resistant rocks, the preservation pattern of the lps is different (fig. 7b). in the south, the lps is mostly preserved as a coherent surface in areas with sedimentary rocks. in areas with crystalline basement, in the south-western part of the study area, for example, the lps is identified as valley shoulders within the major valleys below the ups. in the north, the lps is developed as a coherent surface across a crystalline basement around dove bugt, for example. the lps thus appears to have developed across the entire study area as an extensive plain in coastal areas and as the floor of fluvial valleys incised below the ups in the hinterland. the preservation pattern thus indicates that both the ups and the lps were destroyed quickly when formed across the sedimentary basin, but they have been preserved to a large degree where they formed across basement rocks, or basalts. we thus conclude that, once a large-scale surface has formed across resistant rocks like crystalline basement, it will persist and dominate the topography for a long time and is not easily reshaped (brunsden 1993). the primary destruction of a peneplain thus occurs along the edge of valleys that were formed after uplift of the peneplain. this erosional pattern is observed in other uplifted areas (liu et al. 2019). the effect of glacial erosion on elevated plateaus covered by cold-based ice is limited, while significant erosion occurs in the deeply incised valleys (e.g. sugden 1974; bonow et  al. 2006b; swift et  al. 2008; hall et  al. 2013a; cook et al. 2020). consequently, glacial erosion will lead to an increased relative relief between the comparatively unaffected pre-glacial peneplains and the glacial incision focussed in the valleys (johansson et al. 2001; hall et al. 2013a), whereas isostatic rebound due to fluvial and glacial erosion will increase the overall elevation of uplifted landscapes (medvedev et al. 2013). 6.3 two-stepped surfaces or one arched surface? the geomorphological work by ahlmann (1941) was a major achievement based as it was, on poor maps, limited fieldwork and some aerial photographs. ahlmann (1941) was of the opinion that the landscape in northeast greenland consisted of just one erosion surface (the https://doi.org/10.34194/geusb.v45.5297 http://www.geusbulletin.org bonow & japsen 2021: geus bulletin 45 (1). 5297. https://doi.org/10.34194/geusb.v45.5297 26 of 35 www.geusbul let in.org fig. 20 the incision of the lps into the ups as a valley bench along nordvestfjord (lower and upper planation surface, respectively): a: lps along flyverfjord and nordvestfjord; view north-east across th. sørensen land towards nathorst land (ups) and stauning alper. b: 3d elevation model from aster data in approximately the same area. the lps is developed as valley benches along the main fjords. minor valleys connect the ups with the lps (black arrows). map location in fig. 9. lps a lps lps lps nordvestfjord ups lps ups ups lps ups stauning alper stauning alper flyverfjord flyverfjord ups b elevation (m) 10005000 1500 2000 ups nordvestfjord lps ‘initial topography’ as he referred to it, which generally corresponds to the ups of the current study). he regarded this surface to be arched towards the present-day coast. in contrast, we find that the large-scale landscape in north-east greenland consists of two distinct surfaces with a distinct step between them and not the single arched surface that ahlmann described. where the ups and lps overlap, there is a step between them. this step is not formed by a fault or fracture system (e.g. the post-devonian main fault system); rather, it occurs where the sides of the valleys incised to form the lps. this provides further evidence that the step is erosional. stepped surface landscapes with similar forms are also present along other passive margins in central west greenland (bonow et  al. 2006b), norway (lidmar-bergström et al. 2000; bonow et al. 2003), https://doi.org/10.34194/geusb.v45.5297 http://www.geusbulletin.org bonow & japsen 2021: geus bulletin 45 (1). 5297. https://doi.org/10.34194/geusb.v45.5297 27 of 35 www.geusbul let in.org brazil (peulvast & claudino sales 2004), southern africa (king 1967; green et  al. 2017) and ethiopia (coltorti et al. 2007). 6.4 the domal structure of the ups around stauning alper ahlmann (1941) identified a dome structure for the ‘initial topography’ with the highest elevation occurring around stauning alper. this feature corresponds to the anticlinal structure of the ups identified on profiles for the area in this study (fig. 16). stauning alper exhibits a significant alpine relief and is one of the most dissected areas in north-east greenland. this type of relief is probably not due to a specific rock type because similar crystalline rocks occur in nearby areas at lower elevation where the ups is very well preserved (e.g. renland, milne land; figs 13, 20). it thus seems likely that stauning alper has a tectonic history that deviates from the regional pattern after the formation of the ups because the area experienced more uplift and perhaps, earlier uplift than the surrounding areas. such development would inevitably lead to valley incision and relief rejuvenation, and possibly more erosion by cirque and valley glaciers, which would obliterate the ups. 6.5 comparison with erosion surfaces in west greenland the landscape of central west greenland shares characteristics with that of north-east greenland. a weathered basement surface has been identified in and around the nuussuaq basin (pulvertaft & larsen 2002; bonow 2005). this surface is of cretaceous to paleocene age and is characterised by distinct hills and kaolinitic saprolites (bonow 2005). bonow et al. (2006b) mapped the weathered basement surface south of the nuussuaq basin and demonstrated that it is cut off by a post-basalt planation surface (also labelled ups). the final formation age for the weathered surface in west greenland is similar to the palaeogene, sub-basalt, es in north-east greenland, which indicates that the weathering reflects a large-scale erosion event. the main, post-basalt peneplain (the ups) in central west greenland has similar characteristics to the ups in north-east greenland (bonow et al. 2006a, 2006b). however, whereas the ups in north-east greenland is also preserved in areas with less resistant basalts (e.g. milne land), the equivalent ups in west greenland is mainly preserved in basement rocks and is deeply dissected in the basalt areas on nuussuaq and disko (bonow et al. 2006a). the ups in north-east greenland is located at a significantly higher elevation and is less tilted than that in west greenland. the west greenland ups is also divided into several tectonic blocks that dip in different directions, while the north-east greenland ups is folded only around stauning alper. the high elevation together with less tilting and rotation might explain why the ups is better preserved and thus of a wider extent in northeast greenland. the lps in north-east greenland is present across a significantly wider area than the equivalent surface in west greenland, where it forms only a partially developed surface along a few valleys (bonow et al. 2006b). one possible explanation could be that block movements and tilting have disrupted surface development in west greenland (bonow et  al. 2006b), while the uplifts in north-east greenland apparently have a more regional character. the lps in west greenland was also probably developed across cretaceous sediments and palaeogene basalts in the nuussuaq basin, where the surface was later obliterated by erosion. 6.6 the formation of the lps and the onset of glacial conditions in greenland there is evidence of glacial activity from c. 7 ma off east greenland in the form of ice-rafted debris in ocean drilling program (odp) boreholes, which suggest that full glacial conditions were established in south-east greenland by that time (c. 63°n and 70.5°n; larsen et al. 1994b; jansen et al. 1996; channell et al. 1999). recently, pérez et al. (2018) inferred that cross-shelf glaciation off blosseville kyst and liverpool land began to influence the shelf during the deposition of upper miocene to lower pliocene sediments (7.3–4.9 ma) based on seismic data tied to odp borehole 987, off scoresby sund. for example, they attributed the prograding character of this succession, as well as the shift in depocentre configurations from inner to middle and outer shelf, to ice advancing across the shelf off blosseville kyst. in particular, they attributed outward-bulges in the depositional pattern to active ice streams crossing the continental shelf, in contrast to a previous study that inferred dominantly hemipelagic origin of late miocene deposits off scoresby sund (butt et al. 2001). however, according to pérez et  al. (2018), the two major phases of ice-sheet advance occurred after the miocene – in the early pliocene and around the pliocene–pleistocene transition (at c. 5 ma and between 2.9 and 2.3 ma, respectively). further insight into the glacial history of east greenland may be obtained from the stratigraphic landscape analysis in east greenland (68–78°n) presented here, and by bonow et al. (2014). the formation of the lps by incision below the ups is thought to be controlled by the fluvial system because the lps extends at the same elevation from the valley benches in the interior to the plateau remnants across the sedimentary basins closer to the coast (figs 7b, 9a, 20). this observation implies that full glacial conditions were not established during https://doi.org/10.34194/geusb.v45.5297 http://www.geusbulletin.org bonow & japsen 2021: geus bulletin 45 (1). 5297. https://doi.org/10.34194/geusb.v45.5297 28 of 35 www.geusbul let in.org the formation of the lps. the lps must have formed during the late neogene because it is younger than the ups, which had been formed by the mid-miocene. this interpretation is supported by the integration of stratigraphic landscape analysis with afta data in southern east greenland, which led to the following conclusions (japsen et al. 2014): 1. late miocene uplift (c. 10 ma) led to the onset of incision of the lps below the ups 2. early pliocene uplift (c. 5 ma) led to the incision of valleys and fjords below the lps. this timeframe is further supported by afta data in northeast greenland (japsen et al. in press). the lps was thus established by the fluvial system in the interval between 10 and 5 ma; thus, it is unlikely that full glacial conditions were established in east greenland at this time. we therefore suggest that the fluvial system led to the expansion of the shelf during the late miocene; for example, along blosseville kyst. we note that over-deepened troughs were not observed across the shelf in the corresponding succession (unit 7) on the seismic lines presented by pérez et  al. (2018), whereas major ice streams would have created such troughs. this indicates that glacial action prior to the pliocene was restricted to valley glaciers, which led to deposition of ice-rafted debris offshore, similar to the valley glaciers of greenland reaching the sea today. the prograding character of the upper miocene succession, as well as the shift in depocentre configuration off blosseville kyst at this time, may be the result of glacio-fluvial deposition. the major phases of ice-sheet advance and full glacial conditions thus only occurred after the miocene (solgaard et al. 2013; pérez et al. 2018), in agreement with our conclusion that the final phase of uplift took place in the early pliocene (japsen et al. 2014). 6.7 plateau landscapes are not related to glaciation the presence of mid-miocene volcanics on the ups just south of the study area demonstrates that the ups was formed prior to the onset of large-scale neogene glaciations in greenland (larsen et al. 1994b; storey et al. 2004; bonow et al. 2014). as discussed in the previous section, the lps is the result of late neogene fluvial incision to below the ups. the incision likely occurred prior to uplift in the early pliocene and was followed by the major phases of ice-sheet advance and full glacial conditions. the ups and the lps are therefore not considered glacial surfaces, but their destruction was partly glacial as illustrated by the glacial forms that cut into them. our results agree with the well-established observation that old landscapes, sometime called ‘relict surfaces’ are preserved beneath non-erosive, cold-based ice (kleman 1994; kleman & hättestrand 1999; sugden et  al. 2005; kleman 2008; ebert & hättestrand 2010; hall et al. 2013a, 2013b). however, in terrains of selective linear glacial erosion, wet-based ice behaves differently. in hilly relief terrain, wet-based glaciers exploit and enhance the pre-glacial relief (e.g. sugden 1974) while flat surfaces (e.g. planated surfaces) are not significantly altered (lidmar-bergström 1997; johansson et al. 2001). in north-east greenland, the elevated and flat terrain has facilitated the formation of cold-based ice, even in coastal areas on store koldewey, for example, where the flat basement summits are unscoured (håkansson et al. 2007). alpine relief mainly developed along the coast at stauning alper, for example. similar patterns have been observed in the blosseville kyst region (bonow et al. 2014). braun (2018) argued for the apparent longevity of high-elevation passive margins based on numerical modelling. based on computational experiments, egholm et al. (2017) proposed that plateau landscapes on glaciated margins reflected glacial and periglacial smoothing of an already elevated terrain prior to the inception of glaciers. conclusions from such calculations and modelling are at odds with the results presented here, which document that the extensive plains in east greenland were graded to sea level prior to the glaciations. the suggestion that glaciers produce smooth landscapes at high elevation is also at odds with modern understanding of glacial dynamics (e.g. kleman et  al. 2008; hall & kleman 2014; cook et  al. 2020). finally, the presence of two regional peneplains separated by a major step in the landscape is not explained by the models of egholm et al. (2017). a part of the youngest uplift that produced the second major step in the landscape between the lps and present-day sea level can be explained by glacial processes (medvedev et al. 2013), but isostatic rebound caused by glacial carving cannot produce a two-stepped landscape. we therefore interpret these steps as manifestations of episodes of tectonic uplift that affected north-east greenland during the late neogene. high-level plateaus and deeply incised valleys are characteristic of epcm around the world in both tropical and arctic climates (jessen 1943; king 1967; japsen et  al. 2012a; green et al. 2013, 2018). lidmar-bergström et al. (2000) showed that major landscape features such as high plateaus, a great escarpment and a coastal plain are similar in character in both southern norway and eastern australia. thus, the formation of elevated plateaus along passive continental margins is not necessarily related to glacial action. the evidence presented here for the origin and age of the large-scale landscape in north-east greenland is therefore a typical example of the development of an epcm. https://doi.org/10.34194/geusb.v45.5297 http://www.geusbulletin.org bonow & japsen 2021: geus bulletin 45 (1). 5297. https://doi.org/10.34194/geusb.v45.5297 29 of 35 www.geusbul let in.org 6.8 neogene uplift and the formation of the greenland ice sheet solgaard et  al. (2013) investigated the effects of two phases of neogene uplift in both west and east greenland, on ice sheet initiation using an ice sheet model in combination with a climate model. the uplift scenario tested in that study is similar to that defined by the presence of a landscape in two steps in north-east greenland but put into an absolute timeframe based on afta data from west and east greenland (japsen et al. 2006, 2014). the scenario supposes that the landscape was at a low level prior to the late miocene, after which two phases of kilometre-scale uplift in the late miocene and latest miocene – pliocene (beginning c. 10 and 5 ma, respectively) initiated the formation of the present-day mountains. the landscape model of solgaard et al. (2013) was based on stratigraphic landscape analysis in west and south-east greenland (bonow et al. 2006a, 2006b). the new results from north-east greenland presented in this paper provide further support that the model is valid for greenland in general. the modelling results of solgaard et al. (2013) showed that no ice could build up on the low-lying and almost flat landscape (defined by the ups) prior to the first uplift (c. 10 ma). however, the models demonstrated significant ice expansion in response to increased precipitation and cooling due to the two phases of uplift that raised the ups to its present elevation. large amounts of ice could have formed after the first uplift event (leading to formation of the lps), but the model showed that the ice sheet is sensitive to changes in climate. the second phase of uplift (c. 5 ma, which led to the uplift of the lps) facilitated the ice-sheet build-up and increased the stability of the ice sheet in the models. however, they also suggested a föhn effect that inhibited ice-sheet expansion into the interior of greenland and thus shifted the threshold of formation of the ice sheet towards colder temperatures. under conditions that are colder than the present, the ice could overcome the föhn effect flow into the interior and form a coherent ice sheet. the results thus indicated that the present-day greenland ice sheet is a relict formed under colder conditions. solgaard et al. (2013) concluded that the late cenozoic mountain building in greenland augmented the effects of the climatic deterioration leading to the northern hemisphere glaciations, and that without the second phase of uplift the greenland ice sheet would have been more sensitive to the changes in climate over the past millions of years. the modelling of solgaard et al. (2013) is consistent with the observed climatic variability superimposed on the general cooling trend in the late cenozoic, specifically the presence of ice-rafted debris in late miocene deposits off south-east greenland and evidence for mid-pliocene warmth (larsen et al. 1994a; dowsett et al. 2010). the results are also consistent with the observation that large-scale glaciation in the circum-atlantic region only occurred in the late pliocene, and that the increase in land-ice volume occurred slowly and stepwise over the period from c. 3.5 ma to 2.4 ma (thiede et al. 1998; mudelsee & raymo 2005; pérez et al. 2018). pedersen et al. (2019) used the concept of geophysical relief (small & anderson 1998) to estimate fjord erosion and the subsequent flexural isostatic response to erosional unloading in north-east and north greenland between scoresby sund (70°n) and independence fjord (82°n). they constrained the timing of erosion and isostatic uplift by marine sediments of late pliocene – early pleistocene age that are now exposed on land between 24 and 230 m a.s.l. (feyling-hanssen et al. 1983; funder et al. 2001; bennike et al. 2002, 2010). in this way, pedersen et  al. (2019) found that the independence fjord system formed since c. 2.5 ma. in contrast, their results indicated that fjord formation in the outer parts of scoresby sund commenced before the pleistocene, most likely in the late miocene and continued throughout the pleistocene by progressive, inland fjord formation. the results of pedersen et al. (2019) demonstrate that the inception of the greenland ice sheet began in the central parts of north-east greenland before the pleistocene and spread to north greenland only at the onset of the pleistocene and are thus in accordance with the model of solgaard et al. (2013). 7 conclusions the large-scale landscape of north-east greenland is similar to epcm worldwide, characterised by upland plateaus dissected by deep valleys. the detailed mapping presented in this study reveals a landscape dominated by two regional erosion surfaces, the ups and lps at c. 2 and 1 km a.s.l., respectively, cut by deeply incised valleys. the ups is a major surface at some distance from the coast, up to 200 km wide. in large areas north of 74°n, the ups disappears below the ice sheet. maximum elevation of the ups is found in stauning alper, where the surface is dome shaped. the ups has a low relative relief, a regional extent and cuts across bedrock of variable age and resistance, including crystalline basement and palaeogene basalts that erupted during break-up of the north-east atlantic. this shows that the ups was graded to base level, and the most likely base level was the atlantic ocean after the cessation of the volcanic eruptions. hence, the ups is a post-basalt peneplain. further, the ups must have formed prior to the eruption of the mid-miocene lavas that rest on the ups just south of our study area. a significant change of base level (uplift) terminated the development of the ups, leading to fluvial incision of valleys and to rejuvenation of the relief. the change https://doi.org/10.34194/geusb.v45.5297 http://www.geusbulletin.org bonow & japsen 2021: geus bulletin 45 (1). 5297. https://doi.org/10.34194/geusb.v45.5297 30 of 35 www.geusbul let in.org of base level must have been followed by a period of more stable base-level conditions, because the valleys widened and eventually formed a coastal plain near the base level at that time. consequently, a new regional, low-relief peneplain – the lps – formed. the lps is readily identified in the interior parts of the study area, where it forms fluvial valley benches below the ups at present day elevations of c. 1 km a.s.l. the presence of these well-developed valley benches shows that the lps had formed prior to the onset of late cenozoic glaciations. towards the coast, the valley benches of the ups merge to form a more coherent surface that extends across a 100-km wide zone. where sedimentary rocks crop out, the lps is only defined by flat-topped summits, while in the areas of crystalline basement rocks the lps is more extensive and less dissected. north of 74°n, the lps dominates the ice-free areas of the elevated plateau landscape in near-coastal areas, while the ups further west is covered by the greenland ice sheet. a second phase of base-level lowering (uplift) caused the lps to lose contact with the former base level, and thus the lps was dissected by a new phase of valley incision and relief rejuvenation. the result was a landscape with distinct levels, manifested by a winding escarpment between the ups and lps. this stepped landscape later became glaciated and further eroded. the lps is defined on both sides of the post-devonian main fault system at about the same elevation, and this further supports that the lps was formed by the fluvial system. the escarpment between the ups and the lps does not coincide with the fault system, which appears as a marked scarp in the terrain, defining the western boundary of the sedimentary basin system. each of the uplift events that affected the region after the ups was graded to sea level, were in the order of 1 km and the ups is now at an elevation of c. 2 km a.s.l. the most recent uplift event also resulted in re-exposure of a pre-basalt erosion surface with etch characteristics, resulting from deep weathering in a warm, humid climate. towards the north (germania land and store koldewey) the lps reaches c. 600 m a.s.l., only. the relative denudation chronology established here shows that the topography in north-east greenland was not related to a break-up in the northeast atlantic but developed much later. the denudation chronology for the elevated plateaus in north-east greenland is similar to that previously established for southern east greenland and central west greenland. in these areas, absolute dating from afta documented three events of uplift in the late eocene, late miocene and early pliocene that led to the formation of the ups, the lps and finally to their rise to their present-day elevation. the dramatic change of the east greenland topography since miocene times was an important precondition for the establishment of the greenland ice sheet, and the formation of the lps by the fluvial system shows that full glacial conditions were likely only established after uplift of the lps in the early pliocene. in summary, the ups and the lps provide a visual manifestation of multiple episodes of tectonic uplift that affected northeast greenland in late neogene times. acknowledgements this study benefited from support and interest from many researchers at geus, in particular, asger ken pedersen, james a. chalmers, lotte melchior larsen, stuart watt and troels nielsen. paul f. green commented an early draft of the manuscript. we thank reviewers adrian hall and jean-pierre peulvast for constructive comments, which significantly improved the paper. aster gdem is a product of meti and nasa. additional information funding statement this work was funded by a consortium of oil companies and geus. competing interests the authors declare no competing interests. author contributions jmb: conceptualisation, investigation, methodology, writing – original draft, writing – review & editing. pj: investigation, writing – review & editing. additional files three supplementary files are available alongside this article online, in the geus bulletin data repository: https://doi.org/10.22008/fk2/tmdfp9 references ager, d.v. 1973: the nature of the stratigraphic record, 151 pp. chichester: john wiley & sons. https://doi.org/10.1002/gj.3350290115 ahlmann, h.w. 1919: geomorphological studies in norway. geografiska annaler 1, 3–320. https://doi.org/10.1080/20014422.1919.11880647 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2department of near surface land and marine geology, geological survey of denmark and greenland (geus), aarhus, denmark; 3department of geochemistry, geological survey of denmark and greenland (geus), copenhagen, denmark abstract pesticides and degradation products are a major challenge for groundwater management in europe, and in denmark where drinking water relies entirely on groundwater. to protect drinking water resources, local danish authorities must take groundwater-protective measures in areas designated as sensitive to pollution; however, official zonation for pesticides is lacking. nitrate-sensitive groundwater abstraction areas have been used instead. the goal of our study was to test the appropriateness of this groundwater protection strategy. we used køge municipality (denmark) as a focus area and tested how our findings upscale to the national level. the data for køge municipality included 1070 individual groundwater samples, analysed for at least one of 366 pesticide compounds during the period 2012–2022, which were aggregated at the well-screen level by the median. four pesticide compounds (2,6-dichlorobenzamide (bam), desphenylchloridazon (dpc), n,n-dimethylsulphamide (dms), 1,2,4-triazole) and three pesticide groups (phenoxyalcanoic acids, triazines and dimethachlor and its metabolites) were found with the highest detection frequency in the study area. we found that groundwater pollution with pesticide compounds was not limited to nitrate-sensitive areas in køge municipality or in denmark as a whole. therefore, nitrate-sensitive areas can only be used partially for identifying pesticide-sensitive groundwater abstraction areas. the management implication is that placing protective measures only within nitrate-sensitive areas would be insufficient to fully address the risk of future groundwater pesticide pollution. we identified knowledge gaps and discussed a potential way forward with a more integrated management of groundwater protection in denmark. *correspondence: dv@geus.dk received: 06 jan 2025 revised: 19 may 2025 accepted: 16 jun 2025 published: 28 aug 2025 keywords: pesticides, nitrate-sensitive areas, pesticide sensitive areas, groundwater vulnerability, groundwater protection abbreviations: bam: 2,6-dichlorobenzamide dms: desphenylchloridazon dmsa: dimethylsulfamic acid dpc: desphenylchloridazon eu: european union gko: national groundwater mapping lod: limit of detection m.b.t: metres below terrain nsa: nitrate-sensitive area ow: other well psa: pesticide-sensitive area pw: pollution well rbmp: river basement management plan ww: waterworks well geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: adam hambly (dtu, denmark) reviewed by: niels peter arildskov (watsonc, denmark) and one anonymous reviewer funding: see page 14 competing interests: see page 14 additional files: see page 14 1. introduction groundwater is a critical resource for public water supply in the european union (eu), where it accounts for 65% of the total water abstracted for public water supply (european environment agency 2023). pollution with pesticide residues is a main cause of failure to achieve good chemical status for groundwater in the eu (see ‘regulatory context’ section), only second to nitrate (european environment agency 2018). pesticide substances are a major groundwater management challenge in denmark as well. according to the latest status report (thorling et al. 2024), pesticides or their degradation products were detected in 67.6% of the well screens in the national monitoring network (grumo, n = 1049 for 2020– 2022) and 40.5% of the public waterworks well (ww) screens used for drinking water production (n = 6386 for 2018–2022; thorling et al. 2024). more importantly, the groundwater quality standard of 0.1 µg/l was exceeded at 33.0% of the grumo wells and 10.8% of the ww screens used for drinking water (thorling et al. 2024). this has major implications for the danish drinking water supply, which relies entirely on groundwater. most of it undergoes only simple treatment (aeration and sand filtration), so the overall pesticide status of the treated drinking water is comparable to that of the untreated groundwater (voutchkova et al. 2021). https://doi.org/10.34194/mg1sjj69 https://orcid.org/0000-0003-2840-072x https://orcid.org/0000-0002-1154-3700 https://orcid.org/0000-0001-9067-0005 https://orcid.org/0000-0002-9245-475x mailto:dv@geus.dk https://creativecommons.org/licenses/by/4.0/deed.ast voutchkova et al. 2025: geus bulletin 59. 8390. https://doi.org/10.34194/mg1sjj69 2 of 15 geusbulletin.org danish environmental policy is based on prevention and source protection (pedersen et al. 2016), which is aligned with the eu principles that groundwater quality should be protected by restricting polluting activities in sensitive recharge areas (european environment agency 2023). in denmark, the local authorities (98 municipalities) must take measures to protect the groundwater in areas sensitive to pollution, including pesticides (pedersen et al. 2016). however, national zonation guidelines for pesticide sensitivity are lacking. consequently, some danish municipalities have used nitrate-sensitive areas (nsas, see section 2) as proxies for groundwater abstraction areas sensitive to pesticide leaching. the rationale was that nsas have high infiltration rates, which are assumed to also increase the risk of pesticide leaching (miljøstyrelsen 2000). furthermore, the main source of diffuse groundwater pollution with both nitrate and pesticides is agriculture. there are, however, major differences in their leaching. the most common pesticide compounds in danish groundwater were the persistent transformation products n,n-dimethylsulphamide (dms), desphenylchloridazon (dpc), 4-bis-amido-3,5,6-trichlorobenzenesulphonate (r471811) and 2,6-dichlorobenzamide (bam; thorling et al. 2024). in contrast to nitrate, these compounds do not degrade at the redox front. furthermore, the parent compounds of dms (tolylfluanide and dichlofluanide), and r471811 (chlorothalonil) have also been used as biocides in paint and building materials in urban areas. therefore, we posit that because of differences in the pollution source and geochemical behaviour, the assumption that nsas can be used for groundwater protection from pesticides is problematic. the aim of this research was therefore to test whether the current groundwater management strategy – using nsas as proxies for abstraction areas sensitive to pesticides – is appropriate for protecting danish drinking water resources from pesticides. our working hypothesis was that not only nsas, but also areas outside an nsa can be sensitive to pesticides. we tested our hypothesis for a municipality in denmark with exceptionally high data density and quality, and detailed mapping of nitrate vulnerability and nsas. to determine whether our findings upscale, we performed the analysis at the national level as well. finally, we identified knowledge gaps and discussed different strategies for groundwater protection from pesticides, which is urgently needed in denmark and potentially in other eu countries where drinking water supply depends on groundwater. 2. regulatory context 2.1. legal definition and threshold for pesticides the eu groundwater directive (european commission 2006) defines ‘pesticides’ as active substances in plant protection products and biocidal products, as well as their metabolites, degradation products and reaction products. in many eu countries, there is furthermore a distinction between relevant and non-relevant pesticide metabolites in drinking water (council of the european union 2020), and the eu drinking water thresholds apply only to the former (laabs et al. 2015). a pesticide metabolite is classified as relevant “if there is reason to consider that it has intrinsic properties comparable to those of the parent substance in terms of its pesticide target activity or that either itself or its transformation products generate a health risk for consumers” (council of the european union 2020). the distinction between relevant and non-relevant metabolites is applied to groundwater by many eu member states. however, in denmark and in this study, there is no such distinction – the eu threshold applies to all pesticide compounds in groundwater. laabs et al. (2015) stated that denmark holds a unique position in eu in this regard, but this is in line with the precautionary principle. 2.2. pesticides in eu groundwaters integrated management at the river basin level is key to ensuring the sustainability of groundwater resources in the eu (european environment agency 2023). the groundwater chemical status in eu is assessed as part of the river basement management plans (rbmps), which are the key tool for implementing the water framework directive (european commission 2000). groundwater fails to achieve good chemical status with respect to pesticides, if the eu standard of 0.1 µg/l for individual pesticide compounds or 0.5 µg/l for the sum of pesticide compounds is exceeded (european commission 2006). pesticides are the second-most common reason for failing good chemical status for groundwater. according to the 2nd rbmp (2015–2021), 6.5% by area of the european groundwater bodies failed to achieve good status due to pesticides; moreover, 1.4% by area had an upward concentration trend (european environment agency 2018). the latest assessment (3rd rbmp) showed that 7.5% of the 2050 danish groundwater bodies failed to achieve good status because of pesticides, equivalent to 17% by volume (nilsson et al. 2021). these figures show that groundwater protection from pesticides is a major challenge not only in denmark, but also in the eu. 2.3. nitrate-sensitive areas the danish groundwater abstraction areas designated as nsas, are management areas with a particular sensitivity to nitrate pollution. they are designated under the danish water supply act (miljøministeriet 2022), made public with a ministerial order (miljøministeriet 2023) https://doi.org/10.34194/mg1sjj69 https://geusbulletin.org/ voutchkova et al. 2025: geus bulletin 59. 8390. https://doi.org/10.34194/mg1sjj69 3 of 15 geusbulletin.org and available online (miljøstyrelsen 2023a). nsas are not the nitrate-vulnerable zones from the nitrates directive (council of the european union 1992). the nitrate vulnerable zones are defined as areas of land that drain into polluted waters or waters at risk of pollution and which contribute to nitrate pollution (european commission 2023), while nsas focus only on groundwater in the abstraction areas. a similar distinction between nsas and nitrate-vulnerable zones was made in the uk (cook 1999; osborn & cook 1997). the difference is, however, that denmark is exempt from designating nitrate-vulnerable zones, because it has established and applies nationwide action programmes, according to article 3.5 of the nitrates directive (council of the european union 1992). the main criteria for nsa mapping are the aquifer vulnerability to nitrate and the groundwater recharge (fig. s1; miljøstyrelsen 2000, 2023b; naturstyrelsen 2014). nsas are situated within areas classified as particularly valuable for drinking water abstraction, or within the catchment areas of the ww fields (miljøstyrelsen 2023b). the danish nitrate-vulnerability mapping is based on the characteristics of the aquifer material and the overlaying layers, as well as the groundwater quality (table s1; miljøstyrelsen 2023b). figure 1 illustrates how nitrate vulnerability of the aquifer and the nsa zonation relate. around 17% of denmark (7466 km2) is designated as an nsa. 2.4. pesticide-sensitive areas and pesticide vulnerability we define pesticide-sensitive areas (psas) as those where leaching of pesticides to the groundwater has been observed regardless of the time of application. the pollution with pesticides may result from their application within the entire groundwater catchment area, and not only close to the well head. the historical leaching (their detection in groundwater) is therefore used as a proxy for the areas with inherent pesticide sensitivity. the concentration levels in groundwater, on the other hand, reflect the risk management (e.g. the regulations on dose or time of application) and do not necessarily reflect the inherent pesticide sensitivity. here we do not assess the pesticide sensitivity of large nature areas, where pesticides have not been applied. another inherent limitation is that we cannot test for overlap between nsas and psas where the groundwater is recharged pre1960s, when use of the pesticides in question started. though the pesticide approval procedure has been improved over the years, it still cannot fully prevent leaching of pesticide compounds from approved pesticides, either because the leached degradation products were not identified during the pesticide approval, or because the approval models did not adequately cover real-life conditions. some recent examples are leaching of dms and dmsa from application of cyazofamid (badawi et al. 2024), leaching of tfa from tri-fluorinated pesticides (albers & sültenfuss 2024; johnsen et al. 2024) and leaching of propyzamide in very high concentrations (badawi et al. 2025). the official zonation guidelines (miljøstyrelsen 2023b) do not include national psa designation, but provide a reference (naturstyrelsen 2015), building on the concept by nygaard et al. (2005). this concept uses the soil clay, silt and humus content to evaluate if areas fig. 1 illustration of the principles for designation of a groundwater abstraction area as a nitrate-sensitive area (nsa) in denmark (including aquifer vulnerability and redox state in the aquifer, thickness of reduced clay layer and positive recharge) and potential leaching of persistent pesticide compounds to the aquifer. purple dotted lines indicates potential leaching pathways to the aquifer. gwt: groundwater table (blue dashed lines). fri: first redox interface (red dashed lines). reduced aquifer gwt low nitrate vulnerability >1 5 m reduced aquifer medium nitrate vulnerability 5– 15 m oxic aquifer high nitrate vulnerability fri <5 m >0 mm/y >0 mm/y >0 mm/y reduced clay reduced clay reduced clay pesticides pesticides pesticides (b) nitrate sensitive area (nsa)*(a) outside nsa * only if within areas particularly valuable for drinking water abstraction or within the catchment areas of waterworks wells, otherwise it is outside nsa. https://doi.org/10.34194/mg1sjj69 https://geusbulletin.org/ voutchkova et al. 2025: geus bulletin 59. 8390. https://doi.org/10.34194/mg1sjj69 4 of 15 geusbulletin.org are more sensitive than the preconditions for approving pesticides (nygaard et al. 2005). the concept cannot be applied to soils with >10% clay, wetlands and non-agricultural land (naturstyrelsen 2015). the zonation, therefore, applied to 41.2% of the danish territory, most of which was not sensitive to pesticide leaching. just 0.2% of the land was classified as particularly sensitive, 0.6% as potentially sensitive and 1.2% had a low sensitivity to pesticide leaching (naturstyrelsen 2015). this psa designation had very little practical relevance for the local authorities. pedersen et al. (2016) stated that there are no methods for psa zonation suitable for public administration purposes. according to them, vulnerability to pesticides, should be based on identifying areas with large groundwater recharge, and thus higher risk of pollution, and frequently such areas are also those with sandy soils and sediments, which are also more vulnerable to nitrate leaching (pedersen et al. 2016). 3. methods 3.1. study site køge municipality (255 km2) is situated in the eastern part of denmark on the island of sjælland (fig. 2a). agriculture is the dominant land use (49.4% intensive and 5.4% extensive agriculture); forests cover 20%, while the built-up areas cover 8.4%, and the industry and technological activities take up less than 2% (levin 2019). figure 2b shows the conceptual hydrostratigraphic model of the study area (nw–se cross-section), based on stisen et al. (2020). the pre-quaternary carbonate deposits consist of greensand, bryozoan limestone and chalk from neogene, maastrichtian and campanian; they are depicted collectively as a ‘carbonate aquifer’ in the hydrostratigraphic model of the area (fig. 2b; stisen et al. 2020). the depth to the carbonate aquifer is on average 26 ± 19 m (±1 sd) below terrain (m.b.t) and in the range 0–103 m.b.t (fig. s2a). the overlying quaternary succession (fig. 2) consists primarily of clayey till and sandy meltwater deposits (cowi 2005; jacobsen 2022), formed during the last three glaciations in the pleistocene and impacted by erosion, deposition and deformation. glaciotectonic deformations are expected where the quaternary glacial sequence is thicker, while in the areas with thinner deposits, the clayey till is potentially fractured (cowi 2005; jacobsen 2022). most of the glacial landscape is characterised by till plains with hummocky terrain in places. marginal moraines, erosion valleys, eskers and kames are found as well. there are two buried valleys, eroded into both the glacial sequence and the carbonate aquifer, which were consequently filled with sandy and clayey glacial deposits (sandersen & jørgensen 2016, 2017). the glacial deposit thickness varies from c. 10 m to >100 m, and the thickness of the accumulated clay overlying the carbonate aquifer is on an average 23 ± 19 m and varies from 0 to 102 m (fig. s2b). the nitrate-vulnerability assessment and the nsa mapping (rambøll 2018) were extended with priority action areas which were: (1) within areas of particular drinking water interest and with a high or medium vulnerability, (2) outside forests or natural areas and (3) within the 50-year-catchment zones of the wws (køge kommune 2022). this extended version is used here in this study (fig. 2). within the nsa in the area, the carbonate aquifer depth was on average 15 ± 4 m.b.t. (range 1.3–42.5 m.b.t), while the accumulated clay overlying the carbonate aquifer was on average 13 ± 4 m thick (range 0–34 m). the land use inside and outside nsas can be found in supplementary table s2. 3.2. data 3.2.1. pesticide data the data were downloaded from the nationwide open-access well database, jupiter (hansen & pjetursson 2011) on 10 may 2022. accredited labs upload all chemical analyses of drinking water and groundwater to jupiter. the data extraction was limited geographically to køge municipality, and temporally to samples from the period 2012–2022, and covered 626 different pesticide compounds. the raw data were quality assured (see supplementary text 1) and aggregated at the wellscreen level. only the compounds that were analysed at more than one well screen were retained in the dataset. the cleaned dataset included data from 1070 individual samples analysed for at least one of 366 compounds, representing 452 well screens in 436 wells (some wells have multiple well screens). we focused our analyses on four pesticide compounds and three groups of pesticide compounds, which had the highest detection frequency in the area and were, therefore, of high importance (table 1). bam is a transformation product from the herbicides dichlobenil and chlorthiamide, used in orchards and on paved areas, and from the agricultural fungicide fluopicolide. dpc is a transformation product from the agricultural herbicide chloridazon. dms is a transformation product from the fungicides tolylfluanide and dichlofluanide, used in orchards and production of berries, but dms also leaches from the biocide use of these parent compounds in outdoor paint and wood protection in urban areas (albers et al. 2023). dms is furthermore a degradation product from the agricultural fungicide cyazofamid (badawi et al. 2024). 1,2,4-triazole is a transformation product from a range of triazole-fungicides used in agriculture and as biocides in outdoor https://doi.org/10.34194/mg1sjj69 https://geusbulletin.org/ voutchkova et al. 2025: geus bulletin 59. 8390. https://doi.org/10.34194/mg1sjj69 5 of 15 geusbulletin.org paint. 1,2,4-triazole may furthermore be used as a nitrification inhibitor, though this usage is very limited in denmark. the three compound groups (table  1) were represented by their sum in each sample. the  groups were the agricultural phenoxyalcanoic acids plus their transformation products but excluding the chlorophenols, as they may have other origins, the triazine herbicides and their transformation legend (a) (b) sand clay m –50 –40 –30 –20 –10 0 10 20 30 40 50 60 70 80 carbonate aquifer pre-quaternary clay quaternary sequence nw se 0 2 4 6 8 10km sjælland køge municipality køge jylland 100 km500 denmark 75–80 70–75 65–70 60–65 55–60 50–55 45–50 40–45 35–40 30–35 25–30 20–25 0 2 4 km k øg e b ug t (m.a.s.l.) elevation 5–10 10–15 15–20 køge municipality nsa other wells waterworks wells pollution wells 12 14 16 18 20 22 24 26 fig. 2 overview of the køge municipality study site. a: location of køge municipality in denmark. the groundwater abstraction areas officially designated as nitrate-sensitive areas (nsas) and well locations and types are also shown. elevation in metres above sea level (m.a.s.l). b: conceptual hydrostratigraphic model for the study area. https://doi.org/10.34194/mg1sjj69 https://geusbulletin.org/ voutchkova et al. 2025: geus bulletin 59. 8390. https://doi.org/10.34194/mg1sjj69 6 of 15 geusbulletin.org products that are mostly of agricultural origin and the agricultural herbicide dimethachlor and its transformation products. before aggregation at the well-screen level, the values below the limit of detection (lod) were substituted with 0 µg/l. the aggregation was then based on the median concentration for the period 2012– 2021. for the three groups, first a sum at a sample level was calculated, after which the median was used to aggregate at the well-screen level. in addition, for each well screen a maximum concentration based on all 366 compounds (max366) was calculated. this was done to ensure that all well screens with detections were identified and included in the analyses even when the seven focus parameters were all <lod. 3.2.2. auxiliary data metadata, including well-screen identification number, elevation, well depth, depth to screen top and bottom, x and y coordinates, lithology within the well screen, groundwater body, were also extracted from the jupiter database together with the pesticide data. the nationwide dataset used for the upscaling analysis was also from jupiter database. its preparation was similar to that for køge municipality, methodology details are reported elsewhere (thorling et al. in press). various spatial datasets were also used. the shape files with nsa designations for denmark were downloaded on 24 march 2022 from miljøstyrelsen (2023a). these nsas were supplemented with additional polygons representing priority action areas, provided by køge municipality. the land-use map is a raster in 10 m resolution, which was based on 2018 data (levin 2019). the depth to the carbonate aquifer and the thickness of accumulated clay is from the national water resource model, version 2019 (stisen et al. 2020; ondracek 2022). 3.3. study design 3.3.1. segregation by well type the pesticide data originate from different investigations and monitoring campaigns, reflecting different sampling purposes, and following different analytical programs. based on the purpose of sampling, the well screens were classified as well screens used for: (1) drinking water abstraction (i.e. wws); (2) characterisation of point-source pollution (i.e. pollution wells, pws); (3) other purposes, (i.e. other wells, ows). the ow type included: (1) grumo well screens (thorling et al. 2024); (2) wells used for the national groundwater mapping (gko), including the national nitrate-vulnerability mapping and (3) wells from many other types of investigations, including past wws currently out of use or used for monitoring purposes only. we segregated our assessment by well-screen type (ww, pw, ow), so specific characteristics and biases were accounted for. for example, the pw screens are shallower and shorter than ww screens (table 2). grumo wells (included in the ow type) have short screens (c. 1–2m) and are usually shallower than wws, but deeper than pw screens. the analytical programme is the same for most grumo well screens, which gives a relatively homogeneous subset if only grumo data are used; however, impacts from point-source pollution and urban pollution are underrepresented. there were only five grumo well screens (12.5% of ows) in køge municipality, but this well type has a major importance in the national data set used for upscaling of our findings. we enriched the data analysis by accounting for the specific known biases caused by differing data origin. this topic is further discussed in detail with respect to køge municipality, but the biases and considerations of representativity are also valid at the national scale. table 1 number and percent of well screens with at least one sample with detected or exceeding the groundwater pesticides standard (0.1 µg/l). pesticide compounds cas rn no. of well screens total detected exceeded bam 2,6-dichlorobenzamide 2008-58-4 443 116 (26.2%) 59 (13.3%) dpc desphenylchloridazon 6339-19-1 375 99 (26.4%) 42 (11.2%) dms n,n-dimethylsulphamide 3984-14-3 298 127 (42.6%) 34 (11.4%) 1,2,4-triazole 288-88-0 300 50 (16.7%) 17 (5.7%) phenoxyalcanoic acids (sum) [1] 444 90 (20.3%) 64 (14.4%)[4] triazines (sum) [2] 443 58 (13.1%) 27 (6.1%)[4] dimethachlor + transformation products (sum) [3] 274 52 (19.0%) 27 (9.9%)[4] cas rn is a unique identification number for every chemical substance (registration number assigned by the chemical abstracts service). [1] the group includes 12 compounds, including seven transformation products, see table s3. [2] the group includes 20 compounds, including 12 transformation products, see supplementary table s4. [3] the group includes dimethachlor and seven of its transformation products, see supplementary table s5. [4] the exceedance is calculated based on the sum of the compounds in the specific group for a sample; 0.1 µg/l is used here, even though it is the standard for individual compounds. the 0.5 µg/l standard applies to the sum of all pesticide compounds, so it was not used here. https://doi.org/10.34194/mg1sjj69 https://geusbulletin.org/ voutchkova et al. 2025: geus bulletin 59. 8390. https://doi.org/10.34194/mg1sjj69 7 of 15 geusbulletin.org 3.3.2. statistics and software the pesticide status of each well screen (<lod, detected, exceeding 0.1 µg/l) was mapped and visualised on 1d depth profile, considering both the well type and whether the well is inside or outside nsas. in addition, these sub–data sets were compared based on empirical cumulative distribution functions. we used max. median, which was calculated as the highest value of: (1) the medians of the seven priority pesticides or groups of pesticides (for the period 2012–2022) or (2) the max366 (section 3.2.1). the working hypothesis was tested first by comparing the percentage of well screens with at least one sample with detection (or exceedance of 0.1 µg/l) of the seven pesticides or groups of pesticides inside versus outside nsas. then, we also tested the concentration distributions for well screens inside versus outside nsas for statistical differences. all statistical tests were performed segregated for well type. to make these two comparisons, a spatial join between the locations of the well screens and the nsa polygons was made. the data were not normally distributed, so non-parametric tests were used at the 95% confidence level. the difference in the concentration distributions was tested with kruskal–wallis rank sum test (hollander & wolfe 1973) and post-hoc pairwise wilcoxon rank sum test (hollander & wolfe 1973) with adjusted p-values for multiple comparisons of the ‘fdr’ parameter in r, which controls for false discovery (benjamini & hochberg 1995). summary statistics and statistical tests were handled in the software r v.4.2.1. (r core team 2022). all gis analyses and mapping were done in qgis v. 3.22 (qgis development team 2021). 4. results and discussion 4.1. pesticides in groundwater in the køge municipality 4.1.1. groundwater status: detection and exceedence figure 3 shows the well screens with detections or exceedance of the groundwater quality criterion for pesticides (0.1 µg/l). the well screens were unevenly distributed, as most were located in the south-west, central and north-eastern parts of the municipality. at depth, pesticides were detected down to 39 m.b.t in ww screens, 28 m.b.t in pw screens and 26 m.b.t in ow screens (fig.  3; based on well-screen top). the depth-penetration is most probably underestimated because it is based on the depth to top of the well screen and on the max. median concentration. the priority pesticide compounds (bam, dpc, dms, 1,2,4-triazole) and groups (phenoxyalcanoic acids, triazines, dimethachlor + metabolites) were detected both inside and outside nsas (table 3, fig. 3). the frequency of detection was higher inside than outside nsas for almost all well-screen types and parameters, except for dpc in ww screens, 1,2,4-triazole and dimethachlor + metabolites in pw screens and dpc and dms in ow screens. groundwater status, in relation to drinking water supply and potential compliance issues, is usually assessed based on ww data alone (table 3). all priority compounds and groups were detected in ww screens, both inside and outside nsas, except for 1,2,4-triazole which was not detected outside nsas (detected in 2.9% of ww screens inside nsas). the ww screens had no exceedances for 1,2,4-triazole and the phenoxyalcanoic acids, irrespective of location. the exceedance frequency was larger inside nsas for ww screens, except for dpc which had no exceedances inside nsas, but 4.7% exceedances outside. pesticide detections in wws provide important insights on the state of groundwater because they represent larger volumes of the aquifer due to long screens and large abstraction. this is especially the case in carbonate aquifers, such as the køge municipality, where the median length for ww screens was 16.2 m (table 2). abstracted groundwater, thus, represents a mixture of groundwaters from different depths, including in some cases near-surface-polluted gw mixed with deeper unpolluted or lightly polluted groundwater. however, the frequency of exceedance is lower for ww screens and has less importance here, because of (1) mixing and dilution and (2) closure of polluted wws by the waterworks. the bias in detections and exceedances table 2 summary statistics for length and depth to top of the different types of well screens in køge municipality. well-screen type count (n) na (%) depth to top screen (m.b.t) screen length (m) median ± mad q25–q75 max. median ± mad q25–q75 max. pollution (pw) 307 8.5 10.3 ± 6.3 4.0–17.0 35.7 2.0 ± 0.0 2.0–2.0 19.0 waterworks (ww) 105 8.6 17.4 ± 3.8 14.1–22.8 93.3 16.2 ± 10.2 6.4–31 82.8 other (ow) 40 20 17.4 ± 3.8 10.9–22.5 74.0 6.0 ± 5.0 4.0–41.2 87.3 all 452 9.5 13.0 ± 6.0 7.0–18.0 93.3 2.0 ± 0.0 2.0–6.0 87.3 na: missing depth information for the well screen. mad: median absolute deviation. q25 and q75: the 25th and 75th percentiles. m.b.t: metres to terrain. https://doi.org/10.34194/mg1sjj69 https://geusbulletin.org/ voutchkova et al. 2025: geus bulletin 59. 8390. https://doi.org/10.34194/mg1sjj69 8 of 15 geusbulletin.org at ww screens due to well closures is valid for the pesticides that have been monitored for many years (e.g. bam, phenoxyalcanoic acids, and triazines), and possibly valid for the pesticides added to the analytical programme more recently (e.g. dms, dpc, 1,2,4-triazole). due to the longer well screens and large abstraction volumes, it is assumed that the pesticides detected in ww screens reflect the pesticide pollution in the well catchment area, not only locally in proximity to the well screen. many of the ww screens are outside nsas, but their catchments cover areas both within and outside the nsa (supplementary fig. s3). all these characteristics complicate their use for addressing our hypothesis. therefore, it was necessary to include data from other types of well screens. we assumed that both point sources and diffuse sources could cause pesticide pollution of the aquifer, so the pw data were also included in this assessment. pws are established specifically for characterising pointsource pollution and are usually located near the point source. pollution plumes with pesticides and other micropollutants from point sources in denmark are relatively short (< 250 m), due to dilution from dispersion (bjerg et al. 2021). pws have relatively short intakes (median 2 m, table 2) and only small volumes of water are pumped for the sampling, so it is assumed that the pw data are representative of the leaching and transport of pesticides to the geologic layer where the well screen is located. the pw data contributed to our assessment with information about the spread of pesticides in the near-surface quaternary sediments, overlying the carbonate aquifer. however, some of the pw screens also reached the carbonate aquifer (fig. 3). the pw subset is biased towards high concentrations (see section 4.1.2.), usually for multiple pesticides at a time, high frequency of detections (e.g. up to 48.5% for dms inside nsas, table 3) and high frequency of exceedances (e.g. up to 30.4% for the phenoxyalcanoic acids in nsas, table 3). parent compounds are found more often than in the other well types. exceedances were found in pw screens both inside and outside nsas. the frequency of exceedances was larger inside nsas except for dimethachlor and its metabolites, for which 14.3% of the pw screens outside nsas were with exceedances versus 11% inside nsas (table 3). in this study, we also used data from ows, which is a heterogeneous group, as mentioned in section 3.3.1. the gko wells (n = 10) misrepresent the groundwater status with respect to ‘newly found’ pesticide compounds (e.g. dpc, dms and 1,2,4-triazole) as most of the samples were taken before those compounds were introduced into the mandatory list for monitoring of drinking water (miljøministeriet 2023). all parameters, except the triazines, were detected in ow screens, but ows had no exceedances for dpc, 1,2,4-triazole and the triazines inside or outside nsas. the exceedance frequency for the rest of the parameters was larger inside than outside nsas (table 3). table 3 number of well screens with at least one sample analysed for the seven pesticides and groups of pesticides per well type and percentage of well screens with detected pesticides or with pesticide concentrations exceeding the groundwater quality criterion 0.1 μg/l. pesticide/group of pesticides well type well screens (n) detected (%) exceeded (%) inside nsa outside nsa all inside nsa outside nsa inside nsa outside nsa bam pollution 171 129 300 32.8 25.6 20.5 16.3 waterworks 38 66 104 34.2 10.6 2.6 0.0 other 17 22 39 35.3 4.5 5.9 4.5 dpc pollution 151 105 256 36.4 4.8 23.2 3.8 waterworks 38 64 102 26.3 42.2 0.0 4.7 other 13 4 17 7.7 25.0 0.0 0.0 dms pollution 99 89 188 48.5 32.6 18.2 12.4 waterworks 35 59 94 54.3 33.9 5.7 1.7 other 12 4 16 66.7 75.0 16.7 0.0 1,2,4-triazole pollution 99 91 190 23.2 27.5 9.1 8.8 waterworks 35 59 94 2.9 0.0 0.0 0.0 other 12 4 16 8.3 0.0 0.0 0.0 phenoxyalcanoic acids pollution 171 129 300 38.0 11.6 30.4 8.5 waterworks 38 66 104 13.2 3.0 0.0 0.0 other 17 23 40 11.8 4.3 5.9 0.0 triazines pollution 171 129 300 19.9 16.3 9.4 7.8 waterworks 38 66 104 5.3 1.5 2.6 0.0 other 17 22 39 0.0 0.0 0.0 0.0 dimethachlor + metabolites pollution 100 91 191 16.0 20.9 11.0 14.3 waterworks 26 47 73 34.6 8.5 7.7 0.0 other 7 3 10 42.9 33.3 14.3 0.0 nsa: nitrate-sensitive area. bam: 2,6-dichlorobenzamide. dpc: desphenylchloridazon. dms: n,n-dimethylsulfamide. https://doi.org/10.34194/mg1sjj69 https://geusbulletin.org/ voutchkova et al. 2025: geus bulletin 59. 8390. https://doi.org/10.34194/mg1sjj69 9 of 15 geusbulletin.org 4.1.2. concentration distributions pesticide concentrations were overall higher inside than outside nsas (fig. 4a). this difference was most pronounced for pw screens, which are the dominant type in the municipality (n = 307, 68%). the concentration differences inside or outside nsas for pw screens could be biased because there were more pw wells inside (56%) than outside nsas (44%), which could reflect the focus of the point-source pollution investigations. the concentrations in the wws and ows were three orders of magnitude lower than pws. the pairwise wilcoxon rank sum test showed that the distributions (inside and outside nsas) were significantly different for all three well types (supplementary table s6). these statistical differences could be caused by different hydrogeological conditions inside versus outside nsas. for example, the difference in accumulated clay thickness (supplementary fig. s2b) data: 2012–2022 not detected (<lod) detected (0.01–0.1 μg/l) exceeded (>0.1 μg/l) well-screens: legend wellscreen carbonate aquifer quaternary sand (n = 40) other in (n = 38) out (n = 67) waterworks wells (n = 105) pollution wells (n = 307) in (n = 171) out (n = 136) in (17) out (23) 0 1 2 km pollution wells <lod 0.01–0.1 >0.1 waterworks wells other wells nsa køge 100 80 60 40 20 0 d ep th b el ow g ro un d le ve l ( m ) k øg e b ug t <lod 0.01–0.1 >0.1 <lod 0.01–0.1 >0.1 (a) (b) fig. 3 pesticide status of well screens in køge municipality. a: shown on a map of the municipality along with groundwater abstraction areas officially designated as nitrate-sensitive area (nsas). lod: limit of detection. b: 1d depth profiles of the well screens, ordered by depth to screen top. in: inside nsa, out: outside nsa. n: number of well screens. the entire well-screen length is colour shaded according to the pesticides categories (dark grey, blue, red, as in the legend). some well screens do not reach the carbonate aquifer, in this case there is no colour (white shading). https://doi.org/10.34194/mg1sjj69 https://geusbulletin.org/ voutchkova et al. 2025: geus bulletin 59. 8390. https://doi.org/10.34194/mg1sjj69 10 of 15 geusbulletin.org can delay the transport of pesticides. the accumulated clay layer in nsas is thinner, which could result in a quicker pesticide transport to the carbonate aquifer, if inside the recharge zone. there was also a difference in the land use (supplementary table s2), where the intensive agriculture was 56.2% inside nsas and 46.4 % outside nsas. however, it is not possible to attribute these differences to land-use differences based solely on the data presented here. there could also be other nuances in the hydrogeology. for example, the heterogeneity of glacial deposits is not necessarily reflected in the groundwater models. the usual modelling assumption is that clayey glacial deposits can be represented as homogeneous clay layers in the model, when they could be a mixture of clayey, sandy and silty sediments (in different proportions). moreover, the model resolution is too coarse, and as most operational models are layered (not voxel models), the small-scale heterogeneities in the sediment characteristics cannot be captured. the differences in the distributions for each of the focus pesticides or pesticide groups can be seen in supplementary figs s4 and s5. however, formal statistical tests were not performed due to data limitations (small sub-sets when considering well-screen type, with a high proportion of censored data). 4.2. significance first, the significance of our findings from køge municipality is discussed with respect to the local, regional, national and international scales. then, in section 4.3., 0.01 0.1 1 10 100 0.01 0.1 1 10 100 0.01 0.1 1 10 100 0.00 0.25 0.50 0.75 1.00 pesticides (µg/l) fr ac tio n of w el l− sc re en s location inside nsa outside nsa (a) køge municipality 0. 00 1 0. 01 0. 1 1 10 10 0 10 00 10 00 0 1e +0 5 0. 00 1 0. 01 0. 1 1 10 10 0 10 00 10 00 0 1e +0 5 0. 00 1 0. 01 0. 1 1 10 10 0 10 00 10 00 0 1e +0 5 0.4 0.6 0.8 1.0 pesticides (µg/l) fr ac tio n of w el l− sc re en s location inside nsa outside nsa (b) denmark ow pw ww ow pw ww fig. 4 cumulative distribution of pesticide concentrations by well type inside and outside of a groundwater abstraction area officially designated as a nitrate-sensitive area (nsa). a: køge municipality. b: denmark (thorling et al. in press). note that in b, the y-axis does not start at 0 for visualisation purposes; x-axis is log10-transformed, so the values <lod (0 µg/l) are visualised here as 0.009 µg/l. screen types are as follows: ow: other well; pw: pollution well; ww: waterworks well. https://doi.org/10.34194/mg1sjj69 https://geusbulletin.org/ voutchkova et al. 2025: geus bulletin 59. 8390. https://doi.org/10.34194/mg1sjj69 11 of 15 geusbulletin.org we discuss knowledge gaps (barriers) and possible future directions. 4.2.1. local scale our assessment showed that groundwater pollution with pesticide compounds is widespread in køge municipality (table 3; figs 3, 4). based on the evidence presented here, we confirm our working hypothesis and conclude that while nsas are sensitive to pesticide leaching, they are not the only psas in the municipality. this has major implications for groundwater management locally, as the carbonate aquifer is the main drinking water resource. the local authorities and the public water supplies, which are responsible for protecting the drinking water resources and ensuring safe drinking water, cannot fully address the pesticide issue by applying regulation measures within nsas, as those do not include all psas. the recommended psa zonation method developed by nygaard et al. (2005; see section 2.4.) is not applicable in køge municipality either, as the carbonate aquifer is mostly overlayed by a sequence of clayey layers. 4.2.2. regional scale drinking water resource protection with respect to diffuse pesticide pollution requires trans-boundary cooperation at the regional level within denmark. the 100 year catchment areas of ww screens in køge municipality (supplementary fig. s3) extend outside the administrative border of the municipality. efficient protective measures would target the entire well-catchment zones (i.e. recharge zones) and would not be limited by local administrative borders. to the best of our knowledge there is no existing research focusing on this in denmark. as far as point-source pollution with pesticides is concerned, however, the monitoring and remediation of the polluted sites is handled by the danish regions, even though the impact of the pointsource pollution is usually quite local. 4.2.3. national scale pesticide pollution of groundwater is an issue not only in køge municipality, but also in denmark as a whole (see sections 1 and 2.1). to determine if our findings from køge municipality upscale to the national level, we applied the same methodology on a nationwide data set (thorling et al. in press). the results showed that pesticides were found both inside and outside nsas throughout the entire country, irrespective of well-screen type (table 4). the frequency of detections and exceedances for pw screens inside and outside nsas was very similar, while for wws and ows the frequencies were higher inside than outside nsas. the concentration distributions inside and outside nsas were significantly different for ow and ww screens, but not for pw screens (fig. 4b). this could be explained by the well-screen depth of pws, which are shallower than the other well types. because of the relatively young groundwater in those wells and since they are usually located in proximity to point sources of pollution, they can be expected to be at risk of pesticides pollution irrespective of their location inside or outside an nsa. the ww and ow screens outside nsas should generally have older groundwater ages because of the thicker protective clay layers (according to their definition; fig. 1) delaying the groundwater infiltration (and pesticide transport). however, the reasons for the differences in concentration distributions need further in-depth investigation. table 4 comparison of pesticide status in groundwater for køge municipality and the entire country, inside and outside nitrate-sensitive areas (nsas). max. median well type køge municipality denmark [1] inside nsa outside nsa inside nsa outside nsa well screens (n) pollution 171 136 4943 6093 waterworks 38 67 3238 3450 other 17 23 1980 2607 detected pollution 87% 62% 72% 70% waterworks 79% 57% 51% 27% other 71% 30% 51% 37% exceeded 0.1 µg/l pollution 64% 43% 46% 47% waterworks 8% 8% 12% 5% other 6% 22% 19% 14% median ± mad (µg/l) pollution 0.30 ±0.30 0.06 ±0.06 0.08 ±0.08 0.08 ±0.08 waterworks 0.04 ±0.02 0.01 ±0.01 0.01 ±0.01 0 ±0 other 0.03 ±0.04 0 ±0 0.01 ±0.01 0 ±0 q90 (µg/l) pollution 14.00 2.72 2.20 3.76 waterworks 0.09 0.07 0.12 0.05 other 0.19 0.04 0.23 0.18 summary statistics are based on the max. median concentration. mad: median absolute deviation. [1] data source: thorling et al. in press. https://doi.org/10.34194/mg1sjj69 https://geusbulletin.org/ voutchkova et al. 2025: geus bulletin 59. 8390. https://doi.org/10.34194/mg1sjj69 12 of 15 geusbulletin.org the overall conclusion from the national assessment is that psas are found both inside and outside nsas throughout all of denmark. a major difference is, however, that while the target aquifer in køge municipality is composed of fractured carbonate rocks, there are three other major aquifer types in denmark: (1) quaternary sand, (2) pre-quaternary sand and (3) various older geological units on the island of bornholm (baltic sea). therefore, it can be concluded that irrespective of aquifer type, groundwater protective measures against pesticide pollution placed only within nsas would be insufficient in denmark as a whole. more detailed studies are required to assess the significance of this finding at the local scale throughout denmark. 4.2.4. international scale pesticides are a major groundwater pollution issue in the eu (section 2.2), even though europe’s pesticides regulation is currently one of the most stringent in the world (robinson et al. 2020). foster & custodio (2019) discussed the need for a paradigm shift in water resource management for conserving groundwater in areas under pressure from intensive agriculture, specifically discussing legacy pollutants (nitrate and pesticides). they suggested that groundwater protection requirements need to be incorporated in land-use planning “with zonal restrictions being imposed to reduce risk of pollution according to local conditions” (foster & custodio 2019). however, they did not refer to specific methods for the delineation of groundwater conservation zones, and only stated that those are “well established”. different methods for mapping groundwater vulnerability to pesticides have been applied internationally. pavlis et al. (2010) compared those specific to plant protection products. however, to the best of our knowledge, scientific literature on pesticide-sensitive groundwater abstraction zonation is lacking. it is not possible to directly compare our study with other countries and discuss its international significance in detail, because the peer-reviewed literature on nsas is limited (osborn & cook 1997; cook 1999; foster & custodio 2019). in england, nsas were replaced by the nitrate-vulnerable zones in accordance with the eu nitrate directive in 1996 (foster & custodio 2019). from these literature sources, it was unclear, however, if this only applied to the voluntary or mandatory nature of the measures or also to the spatial zonation. 4.3. future directions 4.3.1. identified knowledge gaps psa zonation is a complex groundwater management issue, which cannot be simplified to the assumption that psas are the same as nsas. we demonstrated this not only locally in the køge municipality, but also at the national scale for denmark. the different behaviour of nitrate and pesticides in the subsurface could explain why this is the case. as mentioned in section 2.3, delineating nsas is based on a vulnerability assessment of the aquifer to nitrate pollution (fig. 1), which is dependent on the redox conditions in the aquifer and the overlaying geological layers (supplementary table s1). details on redox conditions and architecture in denmark are provided in supplementary text 2. the importance of groundwater redox conditions for pesticides is more varied than for nitrate. pesticide compounds are a large and heterogeneous group of organic compounds with very different biogeochemical behaviour with respect to degradation potential and mobility. most pesticide compounds are degraded efficiently under oxic conditions. however, some degradation processes such as dehalogenation may take place under strongly reduced conditions. the most frequently detected pesticide compounds (e.g. dms and dpc) were found both in oxic and reduced groundwater within the køge municipality and throughout denmark, probably because they are persistent with dt50 >100 days (kerle et al. 1996; efsa 2006, 2007) irrespective of redox conditions. for persistent pesticide compounds, there may only be a time-horizon – a lag from the application to their detection in aquifers (fig. 1). there is limited knowledge on the sorption and degradation processes of pesticides in the aquifers, so to better map psas, it is necessary to fill this knowledge gap. the persistence and sorption data are usually from the topsoil, which are not relevant for groundwater and where there is generally lower organic matter content and reactivity. the largest degradation potential for pesticides is in the plough layer, where there is high biological activity and diversity, ensuring microbial degradation of a large proportion of the applied pesticides. the capability of the soil to adsorb pesticide compounds is determined by the soil organic matter, clay minerals and metal oxides (pavlis et al. 2010). the potential to leach to groundwater is also affected by the soil permeability, which is controlled to some extent by the organic matter, and also by soil texture, water fluctuation and water content (pavlis et al. 2010). in soils with high potential for degradation, pesticide compounds could still escape through preferential flow paths like bio-pores and fractures. those are largely unmapped but could explain how pesticides reach the carbonate aquifer in the køge municipality (fig. 3). a discussion on hydrostratigraphic heterogeneity and uncertainty and the scale at which different geochemical and hydrogeological processes can be resolved with respect to pesticides is also needed. the model resolution and type (layered or voxel) would affect the level of detail https://doi.org/10.34194/mg1sjj69 https://geusbulletin.org/ voutchkova et al. 2025: geus bulletin 59. 8390. https://doi.org/10.34194/mg1sjj69 13 of 15 geusbulletin.org to which the inherent sub-surface heterogeneity can be represented. in addition, the representation of geochemical processes would differ depending on the modelling scale and overall framework. 4.3.2. integrated groundwater management there is a need for a nationwide discussion that also includes local groundwater managers, on the practical relevance of ‘one size fits all’ nationwide guidelines for psa zonation. it may be more relevant to adopt a more integrated groundwater protection approach instead of focusing on one pollutant group at a time. however, while such a discussion is necessary, we also urgently need solutions. local managers need to know where to place groundwater-protective measures to safeguard our drinking water resources from future pesticide pollution. pedersen et al. (2016) proposed that mapping vulnerability to pesticides should be based on identifying areas with large groundwater recharge, usually sandy soils and sediments, which are also more vulnerable to nitrate leaching (pedersen et al. 2016). however, we demonstrated here that clayey areas can also be sensitive to pesticide pollution. moreover, sandy (oxic) sediments favour the degradation and sorption of some pesticides. a recommendation needs to be specific to be practically relevant, thus it is necessary to define what is considered ‘large recharge’. should groundwater-protective measures be applied in the entire ww catchment or in the 100 year or 50 year catchment zones? for køge municipality, the modelled 100 year catchment zones of wws (supplementary fig. s3) covered 63% of the municipality and extended beyond its boundaries. it is unclear, if designating all that area as psa would be the optimal solution, due to the lack of uncertainty assessment of the modelled 100 year catchment zones and their relevance when most of the frequently found pesticides have only been used in the past 60 years, but have the potential to persist beyond 100 years. the relevance must also be discussed with respect to pesticide retardation in the subsurface and the target window for groundwater protection. the zonation efficacy should most probably also be discussed with respect to specific protective activities. for example, malaguerra et al. (2012) found that on sjælland (west denmark), wws located in urban areas were more vulnerable to bam and phenoxyalcanoic acids contamination, while non-urban area wells were more often contaminated with bentazon (not included here). urban areas are characterised by a different type of application patterns and source densities in comparison to predominantly agricultural land use. thus, targeting only agricultural areas and different agricultural practices would most likely be insufficient as well. while there has not been an adequate nationwide guidance (a top-down initiative) on how to map psas or how to apply groundwater management measures locally to address the pesticide pollution, a bottom-up initiative for establishing ‘groundwater parks’ (in danish: grundvandsparker) is gaining popularity. groundwater parks are designated areas aimed at groundwater protection, crucial for denmark’s supply of drinking water, but also having a more integrated function involving ecological restoration, afforestation, promoting organic farming, enhancing both nature conservation and climate adaptation efforts (danmarks naturfredningsforening 2021). groundwater parks should typically be established in sensitive groundwater abstraction and recharge areas. in a report for the danish water and wastewater association (danva), refsgaard (2022) outlined some principles for their designation, suggesting using 50 year catchment zones at the 95% confidence level. the ongoing project for establishing groundwater parks in aarhus municipality, however, focuses on 100 year catchment zones of public wws (vpu & aarhus vand 2023). the plan is to establish three groundwater parks in the vicinity of aarhus (c. 330 000 population), converting 4000 ha agricultural land to either nature or forest areas, which would protect 50% of the groundwater recharge areas of the public waterworks (vpu & aarhus vand 2023). the rest of the groundwater recharge is from urbanised areas, where other protection measures are placed by aarhus municipality. the aim of groundwater parks is also to reach other environmental goals for biodiversity and (re)establishment of natural areas (including afforestation and wild self-managed grassland; vpu & aarhus vand 2023). such goals for afforestation are also placed at the national level (‘skovplan’, establishing 250 000 ha of forest). further, the new eu nature restoration law set an overall target that restoration measures should be put in place for at least 20% of the eu’s land area and 20% of its sea areas by 2030; and by 2050, such measures should be in place for all ecosystems that need restoration (directorate general for environment 2024). 5. conclusions the lack of a generally accepted method for pesticide sensitivity mapping in a european context is a challenge for local authorities when they need to implement groundwater protective measures, illustrated here for denmark. we showed that the nsas do not cover all areas of groundwater currently polluted with pesticides. pesticide pollution is widespread both inside and outside of nsas, not only in køge municipality, but throughout all of denmark. we conclude that while nsas are also psas, not all psas are nsas. placing protective measures within nsas, as in some danish https://doi.org/10.34194/mg1sjj69 https://geusbulletin.org/ voutchkova et al. 2025: geus bulletin 59. 8390. https://doi.org/10.34194/mg1sjj69 14 of 15 geusbulletin.org municipalities, will therefore be insufficient to address future pesticide pollution of drinking water. the practical implementation of psas is impeded by knowledge gaps on groundwater-relevant physicochemical properties of approved pesticide compounds. considering the current lack of knowledge, we provide an example of a potential groundwater management alternative, focusing on more integrated approach to safeguarding this sole drinking water resource in denmark. acknowledgements this work was part of research project ‘prioritization – investigation of the use of nitrate sensitive areas (nsa) for measures related to pesticides’, an initiative under the partnership for sustainable water supply, funded by region sjælland and the branch organisation danske vandværker. the authors acknowledge the project workshop participants from dtu-sustain (danish technical university, dtu), region sjælland, køge municipality, the danish epa and danske vandværker, who provided valuable local knowledge, data and discussions. thanks are due also to niels claes (aarhus vand) for critically reading the manuscript and providing an update on groundwater parks initiative. the authors also thank the two reviewers for their comments, which improved the manuscript. additional information funding statement this work was part of research project “prioritization – investigation of the use of nitrate sensitive areas (nsa) for measures related to pesticides,” an initiative under the partnership for sustainable water supply, funded by the region sjælland and the branch-organization danske vandværker. author contributions all authors contributed to the study conception and design. data preparation and formal analysis were performed by ddv. im, lt, and arj provided input on interpretation of results. the first draft of the manuscript was written by dv. im, lt, and arj commented and revised the manuscript. all authors read and approved the final manuscript. competing interests the authors have no relevant financial or non-financial interests to disclose. additional files the following three supplementary files are available at https:doi. org/10.22008/fk2/yxmrlc: supplementary file s1: a .docx file containing tables s1–s6, figs s1–s5, supplementary text 1: data pre-processing and supplementary text 2: redox architecture in denmark. supplementary file s2: aggregated pesticide data set as a csv file supplementary file s3: national (denmark) pesticide data set as a csv file data availability statement the two data sets described in 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https://www.begravededale.dk/pdf_2015/091116_rapport_begravede_dale_bind_1_endelig_udgave_low_res.pdf https://www.begravededale.dk/pdf_2015/091116_rapport_begravede_dale_bind_1_endelig_udgave_low_res.pdf https://www.begravededale.dk/pdf_2015/091116_rapport_begravede_dale_bind_1_endelig_udgave_low_res.pdf https://doi.org/10.34194/geusb.v38.4388 https://doi.org/10.22008/gpub/32631 https://doi.org/10.22008/gpub/38547 https://doi.org/10.22008/gpub/38547 https://doi.org/10.22008/gpub/34784 https://doi.org/10.34194/geusb.v47.6090 https://doi.org/10.34194/geusb.v47.6090 https://deltag.aarhus.dk/node/586/ticket/14930 https://deltag.aarhus.dk/node/586/ticket/14930 the limitations of nitrate-sensitive zoning for groundwater protection from pesticides, denmark 1. introduction 2. regulatory context 2.1. legal definition and threshold for pesticides 2.2. pesticides in eu groundwaters 2.3. nitrate-sensitive areas 2.4. pesticide-sensitive areas and pesticide vulnerability 3. methods 3.1. study site 3.2. data 3.2.1. pesticide data 3.2.2. auxiliary data 3.3. study design 3.3.1. segregation by well type 3.3.2. statistics and software 4. results and discussion 4.1. pesticides in groundwater in the køge municipality 4.1.1. groundwater status: detection and exceedence 4.1.2. concentration distributions 4.2. significance 4.2.1. local scale 4.2.2. regional scale 4.2.3. national scale 4.2.4. international scale 4.3. future directions 4.3.1. identified knowledge gaps 4.3.2. integrated groundwater management 5. conclusions acknowledgements additional information funding statement author contributions competing interests additional files data availability statement references figures fig. 1 illustration of the principles for designation of a groundwater abstraction area as a nitrat fig. 2 overview of the køge municipality study site. a: location of køge municipality in denmark. th fig. 3 pesticide status of well screens in køge municipality. a: shown on a map of the municipality fig. 4 cumulative distribution of pesticide concentrations by well type inside and outside of a gro tables table 1 number and percent of well screens with at least one sample with detected or exceeding the table 2 summary statistics for length and depth to top of the different types of well screens in køg table 3 number of well screens with at least one sample analysed for the seven pesticides and groups table 4 comparison of pesticide status in groundwater for køge municipality and the entire country, geological survey of denmark and greenland bulletin 13, 2007, 45-48 exploration for diamonds in west greenland has experienced a major boost within the last decade following the establishment of world-class diamond mines within the nearby slave province of the canadian arctic. numerous companies have active programmes of diamond exploration and increasingly larger diamonds have been discovered, notably a 2.392 carat dodecahedral stone recovered by the canadian exploration company hudson resources inc. in january 2007. the geological survey of denmark and greenland (geus) is currently carrying out several studies aimed at understanding the petrogenesis of diamondiferous kimberlites in greenland and the physical and chemical properties of their associated mantle source regions (e.g. hutchison 2005; nielsen & jensen 2005). constraint of the mantle geotherm, i.e. the variation of temperature with depth for a particular mantle volume, is an important initial step in assessing the likelihood of such a volume to grow diamonds and hence the diamond potential of associated deep-sourced magmatic rocks occurring at surface. cool geotherms are often present within old cratonic blocks such as west greenland (garde et al. 2000) and provide a good environment for the formation of diamonds (haggerty 1986). this study aims to constrain the mantle geotherm for the southern extent of the north atlantic craton in greenland by applying three-phase geothermobarometry calculations using chemical compositions of clinopyroxene, orthopyroxene and garnet from four-phase kimberlite-hosted lherzolite xenoliths. xenoliths have been sampled from kimberlites from two areas in south-west greenland: midternæs and pyramide fjeld (fig. 1). kimberlites in the pyramidefjeld area principally occur as sheeted sills hosted in the pyramidefjeld granite complex of palaeoproterozoic ketilidian age. in contrast, midternæs kimberlites occur as outcrops within a single, extensive and undulating sill hosted within pre-ketilidian granodioritic gneiss and ketilidian supracrustal rocks. pyramidefjeld kimberlites have been shown to be mesozoic (andrews & emeleus 1971), and work is currently being carried out to further constrain the ages of these and the midternæs kimberlites and also xenoliths using modern methods. no attempt is made herein to provide a correct petrological classification of the rocks hosting the xenoliths; however, the abundance of clinopyroxene reported by andrews & emeleus (1971) suggests that further work may more correctly conclude a classification as ‘orangeite’ after mitchell (1995). notwithstanding this, the term ‘kimberlite’ is employed throughout in order to be consistent with that adopted by previous authors. the precambrian pyramide fjeld granite complex and adjacent archaean grano dioritic gneisses are host to several kimberlite sheets located at various levels between 400 and 900 m elevation (fig. 1a; andrews & emeleus 1971, 1975). kimberlites are mainly found as loose p–t history of kimberlite-hosted garnet lherzolites from south-west greenland mark t. hutchison, louise josefine nielsen and stefan bernstein fig. 1. location of sample sites with reference to context in south-west greenland (index map). red dots indicate location of samples used for pressure and temperature calculation; white dots indicate additional samples. a: sample localities in the pyramidefjeld region. geology simplified from henriksen (1966). b: sample localities in the midternæs region. geology simplified from escher & jensen (1972). 45© geus, 2007. geological survey of denmark and greenland bulletin 13, 45–48. available at: www.geus.dk/publications/bull blocks in scree; however, these are almost always sourced locally from in situ bodies. sheets can often be found deep within overhanging clefts, particularly in granitic walls. the kimberlite bodies are gently dipping, typically 20 degrees, and with a range of strikes. the maximum thickness of sills is approximately 2 m but thickness varies significantly over short distances. in many instances, the occurrence of kimberlite is seen to be controlled locally by structures in the country rocks. field observations of the range of orientations of intrusive bodies do not appear to suggest a particular focal point which could be a likely location for an intrusive centre such as a pipe. this observation is in line with what is seen throughout west greenland where kimberlite emplacement appears as dykes and sills (larsen & rex 1992) rather than the pipes and blows which are common in other world-wide settings. the occurrence of xenoliths amongst pyramidefjeld kimberlites is highly variable with the most xenolith-rich localities being in the vicinity of safirsø (fig. 1a). the majority of xenoliths are dunites with occasional wehrlites and lherzolites (emeleus & andrews 1975). of particular interest from the point of view of thermobarometry is the occurrence of garnet. this is rarely found, even in clinopyroxene-bearing samples, and the two samples chosen for thermobarometry (fig. 1a) represent the majority of the garnet-bearing xenoliths identified within an estimated total population of 75 xenoliths collected. the midternæs kimberlites are hosted in archaean gneisses and proterozoic supracrustal rocks (fig. 1b; andrews & emeleus 1971, 1975). the style of kimberlite emplacement and occurrence of garnet-bearing xenoliths are closely similar to those of pyramidefjeld. contours of elevation be tween outcrops suggest that the kimberlites form parts of a largely contiguous single body dipping at approximately 30 degrees to the west-south-west. individual outcrops as in pyramidefjeld indicate that the body varies in thickness and undulates in response to local structure. the south-western portion of the body which outcrops near the glacier sioralik bræ, is considerably thicker than elsewhere (fig. 2) and in some places is seen to have a true thickness in excess of 4 m. xenoliths are less abundant on average than in pyramidefjeld kimberlites, but a similar variety and proportion of rock types and infrequent occurrence of garnet is observed. the kimberlites from both areas were intruded along zones of platy jointing which likely were caused by degassing of the magma and formed just prior to the kimberlite intrusion. in contrast to some kimberlites in other cratons, very few xenoliths of local, lower crustal rock types have been recognised in the kimberlites from pyramidefjeld and mid ternæs. the intrusions are therefore believed to have been of a non-explosive nature, perhaps because of host-rock rheo l ogy or due to emplacement at relatively deep crustal levels. here we report on calculations of equilibrium pressure and temperature using compositions of three-phase assemblages of garnet, orthopyroxene and clinopyroxene from midternæs and pyramidefjeld mantle xenoliths. measurements polished thin sections of garnet-bearing mantle xenoliths were prepared from fresh samples from both localities. most xenoliths have a coarse granular texture indicative of equilibrium growth amongst clinopyroxene, olivine, orthopyroxene 46 fig. 2. kimberlite sill of approximately 3.5 m thickness intruded within pre-ketilidian gneiss and cross-cutting a vertical gardar age (c. 1200 ma) dolerite dyke. located at midternæs by the glacier sioralik bræ (see fig. 1b). and garnet. typically, triple junctions between mineral grains are well defined. mineral compositions were determined by the jeol 733 electron microprobe at the department of geography and geology, university of copenhagen. ana lyses were conducted using a 15 kv, 15 na and 5 µm beam for the elements si, ti, al, cr, fe, mn, ni, mg, k and na. stand ard isation was achieved against natural and synthetic standards. geothermobarometry estimates of the temperatures and pressures within the earth are essential for the understanding of many geological processes and in particular, in the case of kimberlites, for diamond prospectivity. geothermobarometry is based on the study of mineral assemblages in chemical and physical equilibrium. as mantle xenoliths often retain information about the physical conditions at the time of formation, they are widely used for such estimates. in this study temperature–pressure calculations are based on the two-pyroxene thermometer and the aluminium-in-orthopyroxene / garnet barometer of brey & köhler (1990). results are presented in the context of standard cratonic mantle geotherm models in fig. 3. pressure and temperature estimates from pyramide fjeld range from 909°c and 3.29 gpa to 975°c and 3.50 gpa. these pressures correspond to a depth range of 106–113 km. peak assemblages using cores of touching grains for samples from pyramidefjeld fall on a smooth curve. the mid ternæs sample reflects equilibrium conditions of 1087°c and 3.73 gpa corresponding to 120 km depth and therefore deeper than the pyramidefjeld samples. all values show a similarity with a warm mantle geotherm based on a surface heat flow of 44 mw/m2 after pollack & chapman (1977). however, the location of pressure and temperature points more closely follows the trends in the steady-state geotherm of mckenzie et al. (2005) although at an average temperature elevated by approximately 50°c (fig. 3). furthermore, the midternæs sample may reflect the same type of high-t inflection evident under similar conditions from lesotho kimberlite-hosted xenoliths (finnerty 1989) although further data are required to confirm this inference. discussion and conclusions the apparent coincidence of pressure and temperature values for xenoliths from pyramidefjeld and midternæs along the mckenzie et al. (2005) geotherm suggests that the thermal conditions of the mantle sampled from the two localities separated on the ground by 14 km are largely the same. the mckenzie et al. (2005) model takes account of lower radiogenic heating in the cratonic crust than previously accepted and shows good correlation with pressure and temperature estimates from kimberlite-hosted mantle xenoliths from northern canada and central siberia (mckenzie et al. 2005 and references therein). results from pyramidefjeld and midternæs xenoliths are hence also consistent with kimberlite-hosted xenoliths from elsewhere whilst at the same time elevated temperatures observed in this study suggest that the geotherm was slightly warmer in south-west greenland than in the northern canada and central siberian diamond-bearing mantle. additional samples are required to more closely constrain the geotherm and also to assess the ranges of depths from which xenoliths were sampled at the two locations; however, the greater depth represented so far at midternæs may have significance. midternæs is slightly closer to the central part of the craton, and the greater depths represented in the xenolith suite may thus reflect a thicker cratonic lithospheric root. 47 fig. 3. temperature–pressure diagram showing positions of equilibrium conditions for garnet lherzolite xenoliths in this study in the context of mantle geotherms. diamond–graphite phase boundary after kennedy & kennedy (1976) and steady-state mantle geotherm after mckenzie et al. (2005). mantle geotherms from pollack & chapman (1977) where figures represent surface heat flow for each model in mw/m2. calculated pressures and temperatures are for peak conditions (cores of mineral grains). consequently the present data suggest that the midternæs kimberlite is closer to directly sampling mantle material within the diamond stability field than pyramidefjeld. although xenolith suites may be formed under different conditions compared to diamonds found within the same kimberlites (shee et al. 1982), it appears that midternæs may have a better diamond potential than pyramidefjeld. midternæs kimberlites have not so far been tested for the presence of diamonds, and the apparently shallower-sourced kimberlites from pyramidefjeld have yielded small numbers of diamonds (unpublished company reports collated in jensen et al. 2004). since midternæs contains some of the thickest outcropping kimberlite evident in greenland, the area may merit further attention by diamond prospectors. acknowledgements graham pearson, geoff nowell and nadine wittig (university of durham) are acknowledged for support. mth acknowledges the european community’s 6th framework program for support under a marie curie eif fellowship. disclaimer: this publication reflects the authors’ views, and the european community shall not be held liable for any use of the information contained herein. references andrews, j.r. & emeleus, c.h. 1971: preliminary account of kimberlite intrusions from the frederikshåb district, south-west greenland. rapport grønlands geologiske undersøgelse 31, 26 pp. andrews, j.r. & emeleus, c.h. 1975: structural aspects of kimberlite dyke and sheet intrusion in south-west greenland. physics and chemistry of the earth 9, 43–50. brey, g.p. & köhler, t. 1990: geothermobarometry in four-phase lherzolites ii. new thermobarometers, and practical assessment of existing thermobarometers. journal of petrology 31, 1353–1378. emeleus, c.h. & andrews, j.r. 1975: 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undersøgelse rapport 2005/43, 59 pp. pollack, h.n. & chapman, d.s. 1977: on the regional variation of heat flow, geotherms and lithospheric thickness. tectonophysics 38, 279–296. shee, s.r., gurney, j.j. & robinson, d.n. 1982: two diamond-bearing peridotite xenoliths from the finsch kimberlite, south africa. contributions to mineralogy and petrology 81, 79–87. 48 authors’ addresses m.t.h. & s.b., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: mhutchis@lpl.arizona.edu l.j.n., department of geography and geology, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. research article andrews et al. 2022: geus bulletin 49. 8311. https://doi.org/10.34194/geusb.v49.8311 1 of 20 the sedimentology and depositional environments of the bastians dal and muslingebjerg formations: evidence for the earliest phases of jurassic rifting in north-east greenland steven d. andrews*1 , henrik vosgerau2 , jørgen a. bojesen-koefoed2 1university of the highlands and islands, inverness, uk. 2geological survey of denmark and greenland (geus), copenhagen, denmark abstract the aim of this study is to elucidate the character of the earliest phases of jurassic rifting in northeast greenland. to achieve this, detailed sedimentological analysis and geological mapping were undertaken on the outcrops of central kuhn ø (74°53’55’’n,20°20’56”w). in this region the basement is overlain by the fluvial bastians dal formation (middle jurassic) which is, in turn, overlain by the coal-bearing muslingebjerg formation. a maximum thickness of 140 m is calculated for the bastians dal formation and mapping of stratal geometries demonstrates thinning to both the north and south, confirming that these deposits infill a palaeovalley. predominantly south-westward palaeocurrent orientations are observed and likely reflect the orientation of the palaeovalley (ne–sw). the overlying muslingebjerg formation displays significant lateral variations in thickness as well as facies, thickening from a 5-m-thick coal seam in the north to 50 m in the south. southern outcrops include two intervals of fine-grained sandstones displaying low-angle and trough cross-bedding some of which contain suggestions of tidal bundling. the arrangement of facies suggests that coal formation occurred in both fluvialand shallow-marine (tidal?) environments. coals are similar to those described elsewhere from the muslingebjerg formation and display subtle differences consistent with variable degrees of marine influence. mapping demonstrates the presence of an ne–sw-oriented bounding fault in the south of the region into which the muslingebjerg formation thickens. this likely also controlled the orientation of the underlying ne– sw-aligned palaeovalley and is oblique to the proposed overall n–s orientation of faulting related to rifting through the mid to late jurassic. instead, these alignments resemble those that define pre-jurassic phases of rifting and may therefore indicate a transitional phase of tectonism. faulting on a similar alignment can be traced sw, cutting lindeman fjord and following the valleys east of the a. p. olsen land plateau. *correspondence: steven.andrews914@ gmail.com received: 05 feb 2022 accepted: 16 jul 2022 published: 25 aug 2022 keywords: north-east greenland, jurassic, tectonics, coal, fluvial abbreviations: hi: hydrogen index hst: highstand systems tract oi: oxygen index pi: production index tc: total carbon toc: total organic carbon tmax: temperature of maximum pyrolysate-yield ts: total sulphur tst: transgressive systems tract vr: vitrinite reflectance geus bulletin is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: mette olivarius (geus, denmark). reviewed by: tiffany playter (university of alberta, canada), mihai emilian popa (university of bucharest, romania). funding: see page 19 competing interests: see page 19 additional files: see page 19 introduction alsgaard et al. (2003) discovered fluvial deposits in the central part of the island kuhn ø and described the succession as being 100–150 m thick and infilling a valley system incised into the crystalline basement. they formally defined the succession as a formation and named it after the valley bastians dal, an e–w-trending valley in western central kuhn ø (higgins 2010) where the fluvial deposits were found. their fieldwork was concentrated to the northern part of bastians dal where the fluvial deposits are overlain by a c. 5-m-thick succession of the coal-bearing muslingebjerg formation. overlying shallow-marine sandstones of the pelion formation mark the drowning of the valley fill and the adjacent basement peneplain (alsgaard et al. 2003). the ages of the bastians dal and muslingebjerg formations are not well constrained. the former was assigned a general middle jurassic age by alsgaard et al. (2003) based on its palynological content (fig. 1). fragments of the ammonite arcticoceras cf. ishmae collected from the basal part https://doi.org/10.34194/geusb.v49.8311 https://orcid.org/0000-0002-0418-0809 https://orcid.org/0000-0002-7582-5360 https://orcid.org/0000-0001-5647-2769 mailto:steven.andrews914@gmail.com mailto:steven.andrews914@gmail.com andrews et al. 2022: geus bulletin 49. 8311. https://doi.org/10.34194/geusb.v49.8311 2 of 20 www.geusbul let in.org of the pelion formation in northern wollaston forland (p. alsen, personal communication 2022) indicate that the bastians dal and muslingebjerg formations are no younger than the middle bathonian. however, the boundary with the pelion formation may be largely diachronous (younging towards north) as strata overlying the highest coal bed of the muslingebjerg formation on hochstetter forland contain dinoflagellate cysts suggesting the upper callovian p. athleta ammonite zone (piasecki & stemmerik 2004). the current study presents observations from outcrops covering the entire bastians dal area (74°53’55’’n,20°20’56”w; fig. 2), augmented with photogrammetrical data (fig. 3) and analysis of coal samples and organic residues identified within sandstone samples. these new data shed further light on the ma 155 150 165 160 170 ju ra ss ic m id dl e u pp er l ? b pay pe ug ba onlaps crystalline basement onlaps upper permian mu j l l l l m m m l l m u u u u u tithonian/volgian kimmeridgian oxfordian callovian bathonian bajocian aalenian u no data chronostratigraphy s n wollaston forland – kuhn ø hiatus/condensed alluvial/delta plain – paralic, sand-dominated fluvial and estuarine sandstones, conglomerates shallow-marine sandstones offshore/basinal mudstones, heteroliths coal fig. 1 jurassic stratigraphy of kuhn ø. modified from surlyk et al. (2021). ba: bastians dal formation. mu: muslingebjerg formation. pe: pelion formation. pay: payer dal formation. j: jacobsstigen. ug: ugpik ravine member. b: bernbjerg formation. 2 km fligely fjord bastians dal n baselbjerg quaternary alluvium palaeogene u. jurasic – l. cretaceous. u. jurassic basement m.–u. jurassic basaltic lavas/intrusions wollaston forland gp and younger strata bernbjerg fm payer dal fm and payer dal fm undifferentiated muslingebjerg fm caledonian crystalline basement bastians dal fm stratigraphic boundaries inferred stratigraphic boundaries observed peaks faults logged sections and corresponding figure number c c' 74°30'n clavering ø kuhn ø shannon lindeman fjord hochstetter forland wollaston forland 25 km 19°w20°w21°w 19°w20°w21°w 74°n cretaceous jurassic triassic permian precarboniferous palaeogene quaternary carboniferous greenland a b a b jurassic fault triassic fault a. p. olsen land * * * fig. 4 fig. 7 fig. 8 * fig. 2 location and geology map. (a) overview of the region, based on koch and haller (1971). (b) detailed geological map of the study region, modified from alsgaard et al. (2003, based on koch & haller (1971)) and using data collected during field mapping that formed part of this study. u: upper. m: middle. l: lower. https://doi.org/10.34194/geusb.v49.8311 http://www.geusbulletin.org andrews et al. 2022: geus bulletin 49. 8311. https://doi.org/10.34194/geusb.v49.8311 3 of 20 www.geusbul let in.org architecture, sedimentology and depositional environments of the bastians dal and muslingebjerg formations and the sequence of infilling of the ne–sw-orientated palaeovalley. the orientation of this palaeovalley and associated faulting are thought to relate to a transitional phase of rift orientation between triassic and jurassic rifting events. the new data also reveal that the muslingebjerg formation expands in thickness from the c. 5 m in northern bastians dal to as much as c. 50 in the south. here, coal beds are, in places, intercalated with finegrained sandstones showing signs of tidal influence and resembling the overlying marine pelion formation. at the base of one of the coal beds, oily slicks were observed in out-seeping water, and analysis of the coal samples suggests that they may have acted as a source to the bitumen found within the sandstone pore systems. fig. 3 photogrammetric overview of (a) the northern (bastians dal north) and (b) the southern (bastians dal south) areas of the study region. the locations of key observations and logged sections are indicated, along with the stratigraphic boundaries identified during this study. the numbering (i–ix) of logged sections in a relates to those illustrated in figs 4 and 7. bernbjerg fm ssw a b nne measured sections pelion fm and payer dal fm, undifferentiated top bastians dal fm top muslingebjerg fm top basement top payer dal fm bernbjerg fm s n measured sections observation points top payer dal fm pelion fm and payer dal fm, undifferentiated top muslingebjerg fm top bastians dal fm internal bedding surface top basementfault i ii iii iv v vi viiivii ix fig. 8 https://doi.org/10.34194/geusb.v49.8311 http://www.geusbulletin.org andrews et al. 2022: geus bulletin 49. 8311. https://doi.org/10.34194/geusb.v49.8311 4 of 20 www.geusbul let in.org geological setting the bastians dal and muslingebjerg formations form the basal part of the middle–upper jurassic early syn-rift succession, which is up to 1 km thick in the kuhn ø region. this succession consists of large-scale, backstepping sedimentary units separated by major drowning surfaces, reflecting an overall stepwise northwards transgression. deposition took place on the hanging wall of fault blocks within the wollaston forland basin, which formed a rift-controlled embayment situated on the western margin of the n–s-trending rift complex between greenland and norway. the fault blocks were up to c. 40 km wide fig. 4 summary log through the bastians dal formation (a) and detailed logs of the well exposed sections (b). the general position of the composite section is indicated in fig. 2. positions of the individual detailed logs, numbered i–viii, are provided in fig. 3a. relative positions as indicated on the summary log were calculated using photogrammetry. samples collected from these sections are numbered to the right of the logs. m c si vf s fs m s cs vc s g p co 32 31 40 41 42 43 44 46 47 48 49 50 51 28 29 30 m c si vf s fs m s cs vc s g p co 113 114 115 116 117 131 132 133 134 135 136 137 m c si vf s fs m s cs vc s g p co 70 69 80 81 82 83 102 103 104 105 106 67 68 541236 541237 541238 541213 541214 541217 541218 541219 541220 541221 541222 541223 541224 541215 541216 key horizontal lamination planar cross-bedding trough cross-bedding convolute lamination pebbles coal rootlets coalified wood fine-grained and rooted sandstones & siltstones cross-bedded sandstones trough cross-bedded conglomerates large-scale cross-bedded conglomerates o 10 20 30 40 50 60 70 80 90 100 110 120 130 140 m c si vf s fs m s cs vc s g p co ? ? ? ? ? ? a b i ii iii iv v vi vii viii https://doi.org/10.34194/geusb.v49.8311 http://www.geusbulletin.org andrews et al. 2022: geus bulletin 49. 8311. https://doi.org/10.34194/geusb.v49.8311 5 of 20 www.geusbul let in.org and slightly tilted towards west-south-west. towards the east, elevated fault-block crests formed elongated islands or peninsulas, which broadened and were attached to the mainland to the north. regional sediment transport was overall axial from the north towards the south down a low-gradient basin floor (surlyk 1977, 1990, 1991; surlyk & clemmensen 1983). rifting culminated in the early volgian – early ryazanian with major block faulting, tilting of blocks and the formation of halfgraben (surlyk 1978). vischer (1943) and maync (1947, 1949) were the first to describe the tectonic style and stratigraphy of jurassic tilted fault blocks in the wollaston forland basin. a number of succeeding studies focused on the sedimentary environments, lithostratigraphy and sequence stratigraphy of the early syn-rift succession (e.g. clemmensen & surlyk 1976; petersen et al. 1998; petersen & vosgerau 1999; vosgerau et al. 2000; alsgaard et al. 2003; surlyk 2003; surlyk & korstgård 2013; surlyk et al. 2021). methodology field observations (august 2018) and sedimentary logs (1:40) were augmented with 3d-photogrammetry (for methods, see sørensen & dueholm 2018) to reconstruct the relative position of fragmented outcrops and resolve stratigraphic problems (e.g. the thickness of the bastians dal formation: see fig. 3). coal samples were collected for total carbon (tc), total sulphur (ts), total organic carbon (toc), rock-eval-type pyrolysis, reflected light microscopy and vitrinite reflectance (vr). sandstone samples, which displayed apparent staining, were also collected for toc, rock-eval-type pyrolysis and reflected light microscopy. the methodologies used are summarised here. for full details, see bojesen-koefoed et al. (2020). tc and ts were determined by combustion using a leco cs-200 induction furnace. toc was determined similarly after elimination of mineral-bound carbon by hydrochloric acid treatment, and recalculation based on loss of weight. rock-eval-type screening pyrolysis was done using a hawk (wildcat instruments, humble, tx, usa) calibrated using the ifp160000 pyrolysis standard, ensuring comparability with standard rock-eval data. an in-house standard (marl slate, sunderland quarry, uk) was used for stability control. parameters measured include s1 (hydrocarbons present in the sample mg/g), s2 (pyrolytic hydrocarbons, mg/g), s3 (pyrolytic co2, mg/g) and tmax (temperature of maximum rate of pyrolytic hydrocarbon generation, °c) and the calculated parameters hydrogen index (hi; 100*s2/toc), oxygen index (oi; 100*s3/toc) and production index (pi; s1/(s1+s2)). for details of the rock-eval-type pyrolysis technique and its applications, see espitalié et al. (1985) and bordenave et al. (1993). preparations for reflected light microscopy were made according to standard procedures (taylor et al. 1998) but modified to compensate for variations in settling velocity of particles of different density. the samples were crushed and sieved between 63 μm and 1 mm, and this fraction was embedded in epoxy. blocks were subsequently cut vertically and remounted in epoxy with the cut surface upwards. the mounts were ground and polished to obtain a smooth surface for microscopy. vr (random, oil immersion) was measured using a leica dm4000m reflected light microscope (50 × objective and the diskus fossil system of hilgers technisches buero for recording of measurements). measurements were made at 546 nm (monochromatic light), with the microscope being calibrated against the yag 0.903%ro standard with integrated optical zero. the average vr for each sample was calculated from a population selected from the total reflectance histogram (= %ro). in addition, samples were inspected in reflected white light and fluorescence-inducing blue light using a zeiss microscope. results bastians dal formation the bastians dal formation overlies crystalline basement and is exposed in a series of broken outcrops, which can be traced over several kilometres along an n–s-aligned valley (fig. 3), just beyond the headwaters of bastians dal itself. eight sections were logged in detail (i–viii in figs 3 and 4) through the bastians dal formation. these logs provide the basis of the facies analysis presented here. the facies comprising the bastians dal formation were briefly outlined by alsgaard et al. (2003), who also formally defined the unit. five facies are recognised in the present study: largescale cross-bedded conglomerates, trough cross-bedded conglomerates, cross-bedded sandstones, fine-grained and rooted sandstones and siltstones and thin coals. the latter likely reflect the onset  of conditions, which culminated in the deposition of the thicker coals, which define the overlying muslingebjerg formation. facies: large-scale cross-bedded conglomerates this facies is characterised by large-scale planar crosssets, which form beds up to 1.7 m thick. bed bases are commonly mildly erosive, and relief is often noted across the unit tops. cosets of cross-beds define amalgamated sediment bodies up to 3 m thick and 30 m wide, which contain low-angle internal truncation surfaces (fig. 5a). clasts are dominated by quartzrich lithologies, range from sub-rounded to rounded and reach a maximum clast size of 30 cm, although https://doi.org/10.34194/geusb.v49.8311 http://www.geusbulletin.org andrews et al. 2022: geus bulletin 49. 8311. https://doi.org/10.34194/geusb.v49.8311 6 of 20 www.geusbul let in.org 0.2 m a b c 1 m 1 m fig. 5 photographs illustrating the facies of the bastians dal formation. (a) large-scale cross-bedded conglomerates displaying low-angle truncation surfaces and bar-top relief (c. 69 m, fig. 4). (b) trough cross-bedded conglomerates with sharply erosive bases mantled with pebble (c. 115 m, fig. 4). (c) cross-bedded sandstones. clear alternations of coarser-and finer-grained foresets can be observed, as well as voids, where coalified wood was once present (c. 103.5 m, fig. 4) . https://doi.org/10.34194/geusb.v49.8311 http://www.geusbulletin.org andrews et al. 2022: geus bulletin 49. 8311. https://doi.org/10.34194/geusb.v49.8311 7 of 20 www.geusbul let in.org a maximum clast size of 10 cm is more common within individual beds. coarser material is often concentrated towards the cross-set toes and some normal grading is recognised within individual laminae. pebble-grade clasts are dominant and largely form supported conglomerates. minor very coarse sand to granule material is recorded in some cross-laminae. towards unit tops, tabular-bedded pebbly conglomerates displaying poorly defined horizontal lamination, minor imbrication and occasionally containing large coalified wood fragments (1.5 m long) are also recognised. palaeocurrents measured from both the planar cross-bedding and imbrication within this facies display a strong unidirectional palaeoflow to the sw (fig. 6), which is consistent with the findings of alsgaard et al. (2003). interpretation. the coarse-grained nature of this facies reflects deposition in a very high-energy environment. the strongly unidirectional flow, alongside the largescale cross-bedding and amalgamated nature of the sediment units, suggests deposition as mid-channel bars within a coarse-grained fluvial system (miall 1996). this is consistent with the preserved relief recognised at some unit tops (fig. 5a). individual cross-sets may reflect periods of increased flow within the system, particularly where grading is recognised. more tabular beds towards unit tops record bar-top processes, in some instances providing strand surfaces for woody debris that was washed downstream. facies: trough cross-bedded conglomerates the trough cross-bedded conglomerates contain troughs ranging from 0.2 to 0.5 m deep, which stack to form units up to 1.75 m thick. troughs commonly have erosive bases mantled by coarser pebble lags and show a broad fining-up signature (fig. 5b). clasts up to 30 cm are found in the trough bases but pebble-grade material predominates, often fining-upwards to very coarse sand. clast lithologies are dominantly quartz rich, and rounding ranges from sub-rounded to rounded. imbrication is noted in places, and along with the trough alignment, provide evidence for sw-directed flow. coalified wood fragments are recognised, often concentrated towards the unit tops where the grain size fines towards very coarse sand. interpretation. trough cross-bedding is indicative of migrating sinuous-crested, linked dune forms. these are characteristic of sustained flow within channels and therefore are interpreted to reflect deposition where flow was focused between the mid-channel bars defined by the large-scale cross-bedded conglomerates. the fining-upward signature, often noted in individual troughs, suggests a flashy flow. it was probably during these flash-flood events that the wood fragments were washed into the system before being deposited during waning flow. facies: cross-bedded sandstones cross-bedded sandstones often appear to cap the coarser facies and, furthermore, appear to become more common up-section (fig. 4). planar cross-bedding predominates, forming sets ranging from 0.1 to 0.3 m thick and amalgamated beds up to 1 m thick. the cross-bedding is disrupted to form convolute lamination in places. minor erosion is recognised at bed bases, in some instances defined by pebble lags and coalified wood fragments. coalified wood fragments and associated voids are common throughout this facies. grain size largely varies between very coarse to granule. alternations in these grain sizes can be recognised between individual foresets (fig. 5c), and some laminae display fining-up signatures. compositioanlly the sediments are quartz dominated, but mica booklets up to 4 mm are also observed. this facies includes subordinate horizontally laminated sandstones and structureless sandstones. both have a similar texture and contain ‘floating’ coalfield fragments. interpretation. this facies reflects deposition in much reduced energy conditions, in comparison with the conglomeratic facies. the predominance of planar cross-bedding indicates downstream migration of straight-crested dunes. the limited bed thickness and the association with horizontal lamination (high stage n n = 18 fig. 6 compiled trough cross-bedding, planar cross-bedding and imbrication palaeocurrent data from the bastians dal formation (n = 18). data were collected from thorughout the outcrop area. https://doi.org/10.34194/geusb.v49.8311 http://www.geusbulletin.org andrews et al. 2022: geus bulletin 49. 8311. https://doi.org/10.34194/geusb.v49.8311 8 of 20 www.geusbul let in.org flow) suggest deposition in relatively shallow water. fluctuations in flow intensity, and therefore evidence for flashy flow, are indicated by the alternations in coarser and finer grain sizes between foresets. the fact that this facies caps the coarser-grained facies described earlier suggests deposition within areas of the fluvial system where channels are moving towards abandonment, as well as overbank areas. where this facies begins to dominate the succession in the upper portion of the bastians dal formation, it forms units up to 4.5 m thick. this appears to indicate a transition to lower-energy conditions related to the infilling of topography and the related denudation of the catchment regions. however, basement must still have been exposed locally to source intact mica booklets. where these finer-grained intervals become more dominant in the upper portions of the succession, they are interpreted to record deposition in mobile, high width-to-depth ratio channels. facies: fine-grained and rooted sandstones and siltstones between the coarser-grained facies, much of the bastians dal formation is poorly exposed (fig. 4). in some instances, fine-grained facies are preserved adjacent to the coarser units, and it is suggested that these likely reflect much of what comprises the poorly exposed sections. the sandstones in these exposures are fine to medium grained and form thin beds up to 0.3 m thick. mica and coal fragments are common throughout these sandstones, in which current rippling and small-scale cross-bedding predominate. rooting of bed tops is also common, and in some instances these surfaces are topped by thin (up to 0.2 m) bright, black to blue coals. these thin coals occur in the upper portion of the bastians dal formation and likely herald a transition to the thick coals of the muslingebjerg formation. intercallations of black, papery carbonaceous siltstones up to 0.3 m thick are also noted. interpretation. this facies reflects the lowest energy of deposition recorded in the bastians dal formation. the presence of thin coals and rooted horizons provide evidence for in situ vegetation, leading to the development of organic-rich soils or peats. the presence of thin sands containing current ripples and small-scale cross-bedding likely record overbank flow and related crevasse splay deposition. likewise, the papery siltstones probably reflect inundation of the overbank regions and the subsequent settling out of fines in standing bodies of water. these deposits are typical of overbank facies. stratal geometries and structural implications the bastians dal formation forms a series of broken outcrops that can be traced from the basement to the overlying muslingebjerg formation. eight sections were logged in detail, and the thickness of the intervening non-exposed sections was calculated using photogeology. the coarse-grained elements can be traced across the outcrop area as prominent escarpments for over 500 m (fig. 3a). a maximum thickness of 140 m was calculated for the bastians dal formation, but mapping demonstrated thinning to both the north and the south (fig.  2b), confirming previous suggestions that these deposits infill a palaeovalley. a predominantly south-westward palaeocurrent orientation was recorded throughout the succession (fig. 6), which likely reflects the orientation of the palaeovalley. this alignment is consistent with faulting, which appears to control the thickness of the overlying muslingebjerg formation (described next). thus, faulting may also have exerted a control on the orientation of the palaeovalley when the fluvial deposits of the bastians dal formation were deposited. the fining-upward signature recorded by the occurrence of fine-grained and rooted sandstones and siltstones in the upper part of the formation is consistent with an infilling of topography and the related denudation of the catchment regions. muslingebjerg formation the base of the muslingbjerg formation is not visible where the unit is formally defined on hochstetter forland (surlyk 1977; surlyk et al. 2021). from what is seen in bastians dal, it appears that the basal part of the formation interfingers with fluvial deposits of the bastians dal formation. however, as surlyk (1977) and surlyk et al. (2021) consider the muslingebjerg formation to be characterised by thick coal seams, we regard the base of this unit to lie below the first significant coal seam encountered (1.5 m thick). the upper boundary is placed at the base of the continuous marine sandstones of the pelion formation. in this study, two sections through the muslingebjerg formation were logged in detail: bastians dal north (fig. 7; referred to as the ‘northern section’), where the coals are interbedded with fluvial facies similar to those of the bastians dal formation, and bastians dal south (fig. 8; referred to as the ‘southern section’), where coals are interbedded with the facies described below. three facies dominate in the muslingebjerg formation in the southern section: low-angle cross-bedded sandstones, planar cross-bedded sandstones (with or without bundling) and coal. a single occurrence of trough cross-bedded sandstones is also recorded, and this appears to have closer affinities to the facies found in the bastians dal formation. in the north of the study region, cross-bedded sandstones similar to the underlying bastians dal formation are recorded within the muslingebjerg formation. https://doi.org/10.34194/geusb.v49.8311 http://www.geusbulletin.org andrews et al. 2022: geus bulletin 49. 8311. https://doi.org/10.34194/geusb.v49.8311 9 of 20 www.geusbul let in.org facies: low-angle cross-bedded sandstones the low-angle cross-bedded sandstones form beds up to 2.8 m thick and are characterised by cross-bedding dipping 8–10° to the southwest (fig. 9a). bundling of the cross-laminae is noted within the low-angle cross-bedding, forming laminae bundles 0.1–0.2 m thick. these are sometimes highlighted by the presence of more intensely cemented intervals. small-scale cross-bed sets, 0.15 m thick, are also enclosed within the low-angle cross-sets. well-developed wave ripples are also recorded (fig. 9c). the fine-sand grain size is largely uniform throughout this facies, although some pebble-grade coal fragments are present at some bed bases. only minor bioturbation is recorded in the logged sections but latterally, better defined skolithos and thalassinoides are noted alongside bivalve casts (fig. 9b). m c si vf s fs m s cs vc s g p co (136) (137) (138) (139) (140) (141) (142) (143) 541225 541226 bastians dal north key horizontal lamination planar cross-bedding trough cross-bedding convolute lamination pebbles coal rootlets coalified wood fine-grained and rooted sandstones & siltstones cross-bedded sandstones fig. 7 sedimentological log and accompanying photograph of the thickest coal found in the muslingebjerg formation (ix in fig. 3a). this exposure is found in northern bastians dal. much thinner than its southern counterpart, the associated facies have a greater affinity to the underlying bastians dal formation. location in fig. 2. samples collected from this section are numbered to the right of the log. https://doi.org/10.34194/geusb.v49.8311 http://www.geusbulletin.org andrews et al. 2022: geus bulletin 49. 8311. https://doi.org/10.34194/geusb.v49.8311 10 of 20 www.geusbul let in.org 10 0 m c si vf s fs m s cs vc s g p co 40 41 42 43 44 45 46 47 48 49 50 17 18 19 20 21 22 23 24 25 26 27 28 p el io n fm . m us lin ge bj er g fm . 541231 541230 541229 541228 541232 (from adjacent section) 541227 bastians dal south (muslingebjerg fm.) key horizontal lamination planar cross-bedding trough cross-bedding convolute lamination pebbles coal rootlets coalified wood rooted sandstones trough cross-bedded ssts planar cross-bedded ssts low-angle cross-bedded sandstones fig. 8 sedimentological log through the muslingebjerg formation in southern bastians dal. here, the succession is greatly expanded by the occurrence of shallow-marine sandstones. location and position of the log in figs 2 and 3b. samples collected from these sections are numbered to the right of the logs. https://doi.org/10.34194/geusb.v49.8311 http://www.geusbulletin.org andrews et al. 2022: geus bulletin 49. 8311. https://doi.org/10.34194/geusb.v49.8311 11 of 20 www.geusbul let in.org 0.2 m a b d c 0.2 m 1 m 1 m 0.1 m fig. 9 photographs illustrating the facies of the muslingebjerg formation. (a) low-angle cross-bedded sandstones (c. 18 m, fig. 8). cross-bedding truncated by a horizontal erosion surface towards the top of the section. (b) bivalve moulds and bioturbation within the low-angle cross-bedded sandstones. (c) wave rippling associated with the low-angle cross-bedded sandstones. (d) planar cross-bedded sandstones. bundled foresets are distinctly visible in the lower bed where the bundling is highlighted by coalified wood fragments (c. 21 m, fig. 8). https://doi.org/10.34194/geusb.v49.8311 http://www.geusbulletin.org andrews et al. 2022: geus bulletin 49. 8311. https://doi.org/10.34194/geusb.v49.8311 12 of 20 www.geusbul let in.org interpretation. large-scale low-angle cross-bedding is typical of upper shoreface environments (e.g. ahokas et al. 2014), where it is developed through swash and backwash processes. the assemblage of bioturbation and bivalve casts suggest shoreface development in a marine environment. apparent bundling of foresets may reflect a tidal influence, but could also result from varying energy conditions within the shore zone. small-scale cross-bedding within the larger cross-sets, and the prescence of wave rippling, records the migration of smaller scale bedforms across the shoreface and probably reflects the transition to slightly deeper water. the concentration of coal fragments at bed bases can be interpreted as recording periods of shore-zone reorganisation resulting from low-frequency, high-magnitude storm events. these events will have resulted in erosion of back-beach areas and the liberation of organic material. they may have also generated hummocky cross-stratification in the lower shoreface and offshore transition zone, which would have lain offshore to the south-east. facies: planar cross-bedded sandstones (with or without bundling) this facies comprises fineto medium-grained sandstones displaying planar cross-bedding in beds up to 0.4 m thick. in some examples, the cross-bedding displays well-developed bundling, defined by alternating foresets of coarser and finer grain sizes, as well as variation in the thickness of these couplets. a cyclic inclusion of coaly fragments also occurs in some instances (fig. 9d). regular truncations are noted within the cross-bedding, and the basal portions of the cross-sets often contain concentrations of granule to small pebble-grade material. coalified wood fragments are also common along bed bases. minor trough cross-bedding and horizontal lamination are also observed within this facies. interpretation. the well-developed bundling within the planar cross-bedded sandstones is characteristic of tidal processes. deposition under the influence of tidal currents is recorded by the observed variation in grain size in the forsets, and the bundling of foreset couplets, which reflect longer-term variations (neap-spring) in tidal range. therefore, this facies is interpreted as the deposits of tidal dune fields. occurences of coarser grain sizes and regular inclusion of coalified material suggest a greater link to the fluvial input recorded in the northern section. this is also consistent with the relationship between this facies and the trough cross-bedded sandstones described next. facies: trough cross-bedded sandstones trough cross-bedded sandstones are found as a single interval, c. 1.3 m thick, in the upper sandstone unit of the southern section of the muslingebjerg formation (fig. 8). the sandstones are medium grained and the cross-beds are up to 0.6 m thick, but set thicknesses of 0.1–0.2 m dominate. granules and small pebbles are concentrated along erosive bed bases. interpretation. the facies appears to display affinities with the fluvial deposits of the bastians dal formation. the erosional bed bases and dominance of trough cross-bedding, the latter indicative of sustained flow, along with the absence of marine indicators, favours a similar depositional setting. therefore, the trough cross-bedded sandstones are interpreted as fluvial-channel deposits. this demonstrates that the fluvial system recorded in the facies of the northern section intermittently reached southwards, likely as a result of fluctuations in sea level. facies: coals the coals are best exposed in the northern section where 1.5 and 2.5 m thick coal beds are separated by a thin interval of sandstone (fig. 7). coal beds in the south of the study area, although poorly exposed, appear to reach up to 5 m in thickness. the coals themselves are brittle to friable and bright black to blue. some portions contain alternations of more silty material. the base of the coals is often marked by well-developed root networks, which penetrate underlying sandstones to a depth of up to 0.6 m. in the southern section, an oily slick was observed on water issuing from below one of the coal beds. interpretation. the precursor peats of the coals are interpreted as representing deposition in swampy environments on the fluvial plain (northern section) or coastal plain (southern section). the commonly rooted nature of the underlying sandstones demonstrates autochonous peat formation. the relatively thick coals indicate damming of groundwater, perhaps due to base-level rise caused by a rise in relative sea level. these processes may have resulted in waterlogged, sediment-starved conditions favourable for peat accumulation and preservation. in hochstetter forland, c. 40–50 km north of kuhn ø (fig. 2), liptinite-rich coals of the muslingebjerg formation are present. these have a very good to excellent potential for petroleum generation, which has been interpreted to be governed by marine influence in a coastal-mire depositional setting (bojesen-koefoed et al. 1996; petersen & vosgerau 1999; petersen et al. 1998). the interbedding of shallow-marine sediments with coals in the southern section suggests a similar depositional setting as on hochstetter forland. the coals of the northern sections, which are interbedded with fluvial deposits, were analysed for their source rock potential by petersen et al. (2002) who characterised them as humic coals that https://doi.org/10.34194/geusb.v49.8311 http://www.geusbulletin.org andrews et al. 2022: geus bulletin 49. 8311. https://doi.org/10.34194/geusb.v49.8311 13 of 20 www.geusbul let in.org table 1 results of coal and bitumen analysis. lithology sample number toc tc ts tmax s1 s2 s3 hi oi pi ro n std (wt%) (wt%) (wt%) (°c) (mg hc/g) (mg hc/g) (mg co2/g) (%) sandstone 541228 0.26 0.17 0.04 416 0.00 0.06 0.42 23 162 0.00 sandstone 541229 0.48 0.63 0.06 426 0.02 0.11 0.70 23 146 0.15 coal 541231 51.47 55.99 4.88 418 0.98 51.65 30.46 100 59 0.02 0.36 100 0.025 coal 541232 45.64 49.40 2.45 424 0.73 76.43 10.79 167 24 0.01 0.40 100 0.039 coal 541235 47.65 52.91 3.64 418 0.70 33.89 31.95 71 67 0.02 0.39 100 0.026 toc: total organic carbon (wt%). tc: total carbon (wt%). ts: total sulphur (wt%). tmax: from rock-eval-type pyrolysis, temperature of maximum pyrolysate-yield (°c). s1, s2, s3: parameters from rock-eval-type pyrolysis. hi: hydrogen index, 100 × s2/toc. oi: oxygen index, 100 × s3/toc. pi: production index, s1/(s1 + s2). ro: vitrinite reflectance (%), random mean in oil. n: number of particles measured and used. std: standard deviation of ro population. sample locations in fig. 8. represent only a marginal source-rock type. however, the observation of an oily slick on water originating from the base of one of the coal beds raises the possibility that some of these coals may also have a potential for petroleum generation. the following analysis provides a more detailed investigation of these relationships. coal and bitumen analysis three coal samples were analysed: one from the lower seam (541232) and two from the upper seam (541231 and 541235) – all from the southern exposures of the muslingebjerg formation (results in table 1; locations in fig. 8). all coal samples are thermally immature, showing vr values in the range of 0.36–0.40 %ro and corresponding tmax in the range of 418–424°c. the coals are relatively rich in mineral matter, shown by toc in the range of 45–52 wt%, and based on the difference between tc and toc, significant proportion of the mineral content is carbonate. sulphur contents are rather high (2–5 wt%). the petroleum generation potential of the coal samples is within the range of other samples that have been analysed from the same location (figs 10a, b). all three samples are initially classified as predominantly gasprone, but sample 541232 differs in showing higher hi and lower oi (table 1). using the method of dahl et al. (2004) to assess the properties of the live kerogen fraction, a regression line has been constructed for the ‘lower seam’ of petersen et al. (2002; fig. 10c which demonstrates that this seam indeed has a high proportion of ‘dead carbon’ (25.5%) and that the live kerogen fraction is clearly oil-prone (fig. 10c). sample 541232 falls exactly on this regression line. microscopic inspection of sample 541232 shows a composition dominated by huminite-group macerals, particularly eu-ulminite, but with a fair proportion of resinite present. other liptinite-group macerals are sparsely present (figs 11a, b). fractures within huminite particles frequently show dark discolouration along their margins as well as associated refraction phenomena akin to the well-known ‘newton’s rings’, caused by liquid bitumen filling the fractures (figs 11c, d). two sandstone samples (samples 541228 and 541229) were collected adjacent to the coals (results in table 1; locations in fig. 8). they were suspected of being stained by petroleum, and each yield different results. sample 541228 shows low toc (0.26 wt%), very low pyrolysis yield (s2 = 0.06 mg/g) and a pi of zero. sample 541229 shows somewhat higher toc (0.48 wt%) and a low pyrolysis yield (s2= 0.11 mg/g), but a pi of 0.15, which may suggest light staining. however, the absolute values of s1 and s2 used in the calculation of pi are very small and thus call for caution (fig. 10d). the difference between tc and toc indicates the presence of carbonate. microscopic examination of sample 541229 reveals small patches of bitumen along pores and in pore throats, and pores lined with a thin carbonate cement (figs 12a, c). bitumen shows a dull, brownish fluorescence, whereas the carbonate cement appears yellow (figs 12b, d). interpretation the muslingebjerg formation in bastians dal represents the transgression of an infilling palaeovalley. hence, coals in the southern outcrops and the lower part of the succession may be more prone to marine influence (resinite enriched) than coals which formed updip, and higher in the succession (humic), and thus show higher petroleum potential (petersen et al. 1998). elements of marine influence are also hinted at by the slightly higher than expected levels of sulphur observed, reflecting the relatively sulphate-enriched character of marine water. thermal maturity is low, and clearly below the threshold for petroleum generation. however, the presence of bitumen-filled fractures in sample 541232 (lower seam) is still not surprising, since some resinite is known to generate liquid petroleum at a very early stage in the maturation process (e.g. snowdon & powell 1982; khorasani & murchison 1988; snowdon 1991). https://doi.org/10.34194/geusb.v49.8311 http://www.geusbulletin.org andrews et al. 2022: geus bulletin 49. 8311. https://doi.org/10.34194/geusb.v49.8311 14 of 20 www.geusbul let in.org 400 450 500 550 tmax (°c) 0 100 200 300 400 500 600 700 800 900 1000 h yd ro ge n in de x (h i) 541231 541232 541235 muslingebjerg fm. coals bastians dal, profile 1, upper seam (petersen et al. 2002) bastians dal, profile 1, middle seam (petersen et al. 2002) bastians dal, profile 1, lower seam (petersen et al. 2002) bastians dal, profile 1 (petersen et al. 2002) bastians dal (petersen et al. 2002) bastians dal, profile 2, single seam (petersen et al. 2002) bastians dal, this study 0.1 1 10 100 toc (%) 0.1 1 10 100 s 2 (m g h c /g ro ck ) excellent excellent good g ood fair poor poor 0 20 40 60 80 100 toc (%) 0 100 200 300 s 2 (m g h c /g ro ck ) 541231 541232 541235 muslingebjerg fm. coals, lower seam (pink symbols): dead carbon = 25.5% hilive = 343 380 390 400 410 420 430 440 450 460 470 480 490 500 510 520 tmax (°c) 0.0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1.0 p ro du ct io n in de x (p i) (s 1/ (s 1+ s 2) ) 54 12 28 54 12 29 coal samples sandstone samples immature oil condensate, wet gas dry gas staining or contamination low-level conversion a b c d 541232 541231 541235 fig. 10 rock-eval-type screening pyrolysis data for samples from bastians dal. for comparison, data from the present study are plotted together with data on coal samples from the same location, published by petersen et al. (2002). (a) standard-plot of tmax vs. hydrogen index (hi). samples from the present study all fall within the range of previously published samples. sample 541232 shows somewhat higher petroleum potential than samples 541231 and 541235. (b) standard-plot of toc vs. s2. samples from the present study all fall within the range of previously published samples. (c) linear plot of toc vs. s2. following the method of dahl et al. (2004), a regression line has been constructed for the lower, more prolific seam, suggesting high levels of inert carbon, but also high petroleum potential of the live kerogen fraction (hilive). sample 541232 falls on the regression line. (d) standard plot of tmax vs. production index (pi). only samples from the present study are shown. sandstone sample 541229 appears lightly stained, whereas 541228 does not. https://doi.org/10.34194/geusb.v49.8311 http://www.geusbulletin.org andrews et al. 2022: geus bulletin 49. 8311. https://doi.org/10.34194/geusb.v49.8311 15 of 20 www.geusbul let in.org petersen et al. (2002) note that in cases where the proportion of liptinite in coals from bastians dal exceeds a few per cent, the excess liptinite is resinite. since resinite is the main liptinite-group maceral observed in sample 541232, we assume that the liquid bitumen observed represents products of early generation from resinite. highly oil-prone resinite-rich coals (payer dal, on kuhn ø and hochstetter forland, north of kuhn ø) are known from the muslingebjerg formation (bojesen-koefoed et al. 1996; petersen et al. 1998, 2002). such coals are even found as erratic blocks on the coasts of norway (horn 1931; bojesen-koefoed et al. 1996; petersen et al. 2013). these coals are paralic, representing selective concentration of liptinite-group macerals, particularly resinite. humic coals are also common in the muslingebjerg formation, such as those found in the northern outcrops visited in this study, and reported by petersen et al. (2002), which are not conspicuously oil-prone (petersen et al. 1998, 2002; bojesen-koefoed et al. 2012). the findings of the present study are in line with data published previously by petersen et al. (2002). the close similarity between sample 541232 and samples of the ‘lower seam’ of petersen et al. (2002) is striking. although the absolute values of s1 and s2 yielded by the sandstone sample 541229 are indeed low, the presence of light staining is confirmed by microscopy, with the dull fluorescence probably resulting from oxidation. given the low level of thermal maturity of the succession in general, this result is somewhat surprising. although generation of petroleum at low levels of thermal maturity is demonstrated by the coal sample 541232, expulsion of petroleum products, that is, secondary migration, requires prior saturation of the pores of the source rock. this will not take place unless the concentrations of resinite are much higher than those observed here. the presence of staining, even at the trace level observed here, could indicate that intervals with higher resinite concentrations are present somewhere in the succession, and that such intervals are able to expel small amounts of petroleum products even at fig. 11 photomicrographs of coal sample 541232. width of field approximately 300 µm. (a) huminite particle (grey) enclosing lump of resinite (bright pinkish yellow). (b) field as a. fluorescence-inducing blue light. (c) huminite particle with minute fractures showing discolouration along the rims due to liquid bitumen. refraction anomalies akin to ‘newton’s rings’ are caused by bitumen from the fractures dissolving into the immersion oil used for microscopy. (d) same field as c. fluorescence-inducing blue light. note the presence of fluorescence of the petroleum-stained rims of the fractures and blooming phenomena at refraction anomalies. a b c d https://doi.org/10.34194/geusb.v49.8311 http://www.geusbulletin.org andrews et al. 2022: geus bulletin 49. 8311. https://doi.org/10.34194/geusb.v49.8311 16 of 20 www.geusbul let in.org low levels of thermal maturity. at present, little evidence is available to substantiate this, although such prolific resinite-rich deposits are known from payer dal (kuhn ø) and hochstetter forland. stratal geometries and structural implications the muslingebjerg formation thickens rapidly from north to south, across the exposures of central kuhn ø. in the north, the muslingebjerg formation is c. 5 m thick. here it comprises two coal beds separated by 0.5 m of muddy siltstones and a thin coarse sandstone bed (fig. 6), which closely resembles the facies of the underlying bastians dal formation. in the south, a more complex and expanded (c. 50 m thick) succession is recorded (fig. 8). this consists of three finer-grained and more poorly exposed units, divided by two thick sandstone successions. the finer-grained units appear to be dominated by coaly material, but the poor exposure does not allow more detail to be gained. the coals are best exposed immediately above the sandstone units, which often contain extensive rooting below the contact. in the southern outcrops, the lowermost fine-grained interval (13 m thick) contains thin intercalations of coarse-grained sandstones, similar to the underlying bastians dal fomation facies. as described earlier, it seems likely that peat accumulation was initiated due to a rising water table linked to a rise in relative sea level. a continued rise in relative sea level led to the transgression of the swampy, peat-forming environment, and eventually the depostion of a shallow marine succession, comprising shoreface and tidal dune-field elements, which display an overall coarsening-upward motif. up-dip, in the northern outcrop region, fluvial deposition continued, and at times the fluvial influence can be recognised to extend southwards (trough cross-bedded sandstone facies). coal, with associated rooting, caps the lower sandstone unit and likely reflects a basinward shift in facies, resulting from falling relative sea level. however, the continued accumulation of coal suggests elevated water tables and therefore a return to transgressive conditions. the capping of this second coal with a coarsening-upward package of shallow-marine sandstones suggests the eventual flooding of the peat-forming environment before relative sea level stabilised, as described from the lowermost coal-sandstone couplet. this pattern is repeated twice more, before being capped by the thick shallow-marine pelion formation. the marked thickening in the muslingebjerg formation teminates abruptly against a ne–sw-aligned fault in the south of the study area (fig. 2). this previously unknown structure defines the southward extent of the muslingebjerg formation, with only a much reduced thickness of the bastians dal formation continuing south across the fault. the thick development of the muslingebjerg formation and its intercalation of marine a b c d fig. 12 photomicrographs of sandstone sample 541229. width of field approximately 300 µm. (a and c) rounded quartz grains, pore surfaces lined with patchy, thin coatings of bitumen and carbonate cement. (b) same field as a. fluorescence-inducing blue light. (d) same field as c. fluorescence-inducing blue light. pore-throats show accumulations of clay and dull brownish-fluorescing bitumen and other organic particles. pores are lined by a thin carbonate cement (yellow fluorescence). https://doi.org/10.34194/geusb.v49.8311 http://www.geusbulletin.org andrews et al. 2022: geus bulletin 49. 8311. https://doi.org/10.34194/geusb.v49.8311 17 of 20 www.geusbul let in.org sandstones to the south reflect enhanced accommodation space available for deposition in this area. to the north, the muslingebjerg formation is considerably thinner and the coals are intercalated with fluvial sediments. the variation in stratal thickness and facies was probably governed by syn-sedimentary subsidence along the fault, which limits the extent of the muslingebjerg formation to the south. placed into a wider regional sequence stratigraphic framework, in the broadest sense, the muslingebjerg formation can be viewed as part of a transgressive sequence tract (tst), where the boundary to the overlying pelion formation represents regional marine flooding. this of course can be broken down into considerably more detail. petersen et al. (2002) considered each coal-sandstone couplet to represent a sequence, with the rooted surface found at the sandstone top defining the sequence boundary (reflecting a basinward shift in facies). in this interpretation, the coal formed during the tst and the sandstone formed during the highstand systems tract (hst). even higher-order units were suggested by petersen et al. (1998), where dulling-upward cycles were recognised within the coals. coal analysis of comparable detail was not possible during this study. an alternative interpretation could be considered where the lowermost and poorly exposed section of the muslingebjerg formation in the southern outcrops would represent a tst, with the upper portion, including the two shallow-marine sandstone units, recording forestepping parasequences of the hst (e.g. holz et al. 2002; ketzer et al. 2003). if this were the case, the sandstone and coal couplets would be genetically linked to the overlying pelion formation. geological structure fault control on the distribution and thickness of the muslingebjerg formation provides evidence for the earliest stages of jurassic rifting in the region. north of kong oscar fjord (c. 72°n), the early jurassic is characterised by uplift and erosion. this is replaced by rift-related tilting of fault blocks and onlap of the pre-jurassic basement from the bathonian with rift climax reached during the latest jurassic (surlyk & ineson 2003). jurassic rifting in this region is largely characterised by n–s fault orientations (surlyk 2003; guarnieri et al. 2017). this is a distinct change from the ne–sw-oriented faulting that dominates the triassic (guarnieri et al. 2017; andrews et al. 2021). the bastians dal and muslingebjerg formations are constrained as no younger than the middle bathonian (see discussion in the introduction section) and therefore provide an important insight into the nature of the earliest phases of rifting in the region. the earliest deposits, the bastian dal formation, infill an ne–sw-aligned palaeovalley. this is oblique to the proposed n–s orientation of faulting related to rifting through the mid to late jurassic (surlyk 2003; guarnieri et al. 2017). the exploitation of pre-existing basement weaknesses related to earlier phases of rifting could explain palaeovalley development on such an alignment. the discovery of an ne–sw-oriented fault in the south, into which the muslingebjerg formation thickens, confirms that active rifting on ne–sw-oriented faults occurred at this time. post pelion formation, reverse movement also appears to have occurred before the fault became inactive, as evidenced by displacements on the pelion formation and the uninterrupted nature of the overlying bernbjerg formation (fig. 3b). it seems likely that the alignment of the bastians dal palaeovalley was therefore, to some extent, also controlled by the fault activity recognised to control the thickness of the muslingebjerg formation. the ne–sw alignment of the palaeovalley and the faulting that controlled the distribution of the muslingebjerg formation are similar to pre-jurassic and triassic phases of rifting (guarnieri et al. 2017; andrews et al. 2021). furthermore, faulting on this alignment can be traced sw to a fault, suggested to be of triassic age (figs 2 and 9 in guarnieri et al. 2017), which cuts lindeman fjord and follows the valleys east of the a. p. olsen land plateau, suggesting more than just a localised occurrence. if the earliest phases of middle jurassic (bathonian) rifting followed the structural trends established during the triassic, it would suggest that it was during this period that a shift in stress fields occurred. constraining the timing of this shift may help us elucidate the controls on these changes. this discovery further implies that the proposed triassic structural grain is consistent northwards from jameson land (c. 70–72°n) where it was first recognised and therefore is likely to have also exerted a major control on facies distribution in more northern regions. synthesis the bastians dal formation (140 m) records the deposits of a high-energy gravel-bed river that flowed through an ne–sw-oriented palaeovalley (fig. 13). large-scale gravel bars were separated by channels, along which migrated sinuous-crested dunes. the outcrop available does not allow for the confident identification of the fluvial planform. recent studies (e.g. hartley et al. 2015; swan et al. 2018) have demonstrated that great caution should be taken when attempting to identify planform from limited two-dimensional outcrops. flow was likely flashy, and during flood events overbank areas accumulated straight-crested dunes and finer-grained deposits, some of which supported vegetation. thin https://doi.org/10.34194/geusb.v49.8311 http://www.geusbulletin.org andrews et al. 2022: geus bulletin 49. 8311. https://doi.org/10.34194/geusb.v49.8311 18 of 20 www.geusbul let in.org coals are recognised towards the top of the bastians dal formation, suggesting a transitional contact with the overlying muslingebjerg formation, the base of which is marked by the first thick coal. the muslingebjerg formation displays significant lateral variations in thickness as well as facies, thickening from a c. 5 m thick coal seam in the north to c. 50 m in the south (fig. 13). the coals of the northern region of the study area contain minor intercalations of fluvial deposits, and these are replaced southwards with shoreface, and tidal facies associated with shallow-marine conditions (fig. 13). these changing environments are also recorded in the coals, with a more humic character recorded in the north (petersen et al. 2002), and, initially, more resinite-enriched coals in the south. placed in a regional sequence stratigraphic context, the muslingebjerg formation appears to form part of a tst, where the overlying boundary to the pelion formation ? + + + + + + + + + + + + + + + + + + + + ++ + + + + + + + + + + + + + + + + + + + + + ++ + a. bastians dal formation b. muslingebjerg formation north fig. 13 synthesis of sedimentological observations and stratigraphic geometries recorded in the bastians dal region. (a) infilling valley that existed during the deposition of the bastians dal formation. a gravel-bed fluvial system dominated deposition with finer-grained overbank facies becoming more common over time and including minor coal developments. the sw alignment of the valley may have been controlled by incipient faulting, which later also controlled the distribution of facies and stratal geometries within the muslingebjerg formation. palaeoflow of the fluvial system is indicated by the arrow. (b) interaction of environments recorded during the deposition of the muslingebjerg formation. rising relative sea level and reduced topography in the hinterland led to widespread coal formation (dark grey). these were overlain by shallow-marine shoreface and tidal (arrowed circulation: dashed line) sandstones as sea level continued to rise. with continued regional transgression, the muslingebjerg formation gave way to the shallow-marine sandstones of the pelion formation (not shown). palaeoflow of the fluvial system is indicated by the thin arrow (continuous line). as is clear from the stratal geometries described in this study, subsidence during this period was fault controlled. legend for sedimentary structures in figs 4 and 8. https://doi.org/10.34194/geusb.v49.8311 http://www.geusbulletin.org andrews et al. 2022: geus bulletin 49. 8311. https://doi.org/10.34194/geusb.v49.8311 19 of 20 www.geusbul let in.org represents regional marine flooding. petersen et al. (2002) interpreted the coal-sandstone couplets of the muslingebjerg formation as higher-order sequences. alternatively, these could be interpreted as sandstone-coal parasequences, comprising forestepping marine sandstones and associated coals deposited during the hst (e.g. holz et al. 2002; ketzer et al. 2003) and therefore genetically linked to the pelion formation. three coal samples and two sandstone samples were analysed for petroleum potential and oil staining. the samples are thermally immature, and none of the coal samples can be classified as oil-prone. however, recalculations of the properties of the live kerogen fraction based on the method of dahl et al. (2004) suggest that sample 541232 is oil-prone, but that this feature is masked by a high concentration of ‘dead’ carbon. this is confirmed by the demonstration of bitumen-filled fractures in huminite macerals, probably caused by early generation from resinite – the more abundant among the generally sparse representatives of the liptinite maceral group. one of the sandstones is lightly stained by bitumen, indicated by both pyrolysis and microscopy. this was unexpected given the low level of thermal maturity of the coals and the associated problems of overcoming the expulsion threshold. staining could possibly indicate the existence of intervals of higher resinite concentrations within the succession, where this threshold can be locally overcome. the thickening of the muslingebjerg formation is demonstrated to have been controlled by an ne– sw-oriented bounding fault in the south of the region. it seems likely that this also controlled the orientation of the underlying ne–sw-aligned palaeovalley. this is oblique to the proposed overall n–s orientation of faulting related to rifting through the middle to late jurassic (surlyk 2003; guarnieri et al. 2017). instead, these alignments resemble those of pre-jurassic phases of rifting (guarnieri et al. 2017; andrews et al. 2021) and may therefore indicate a transitional phase of tectonism. faulting on this alignment can be traced sw, aligning with the proposed triassic fault of guarnieri et al. (2017), which cuts lindeman fjord and follows the valleys east of the a. p. olsen land plateau. conclusions the relationship between n–s oriented jurassic rifting and earlier ne–sw-oriented triassic rifting in east greenland has received little attention. deposits of middle jurassic age from central kuhn ø appear to demonstrate a continuity with triassic structural trends, suggesting that a transition occurred during the early to middle jurassic. the south-westward flowing fluvial facies of the bastians dal formation infill a ne–sw-oriented palaeovalley, the orientation of which may reflect incipient rifting. the overlying muslingebjerg formation is defined by the occurrence of coal, which is intercalated with both fluvial (in the north) and shallow marine facies (in the south). the coals are similar to those described elsewhere from the muslingebjerg formation but display subtle differences consistent with variable degrees of marine influence. a distinct nw–se thickening of the muslingebjerg formation is demonstrated to be controlled by a ne–sw-oriented fault. the ne–sw orientation of palaeotopography, palaeocurrents and faulting identified in central kuhn ø is consistent with the orientation of triassic rifting and can be traced sw to faulting that cuts across lindeman fjord and follows the valleys east of the a. p. olsen land plateau. to understand the geographic and temporal transition between the triassic rifting and n–s-oriented jurassic rifting, a detailed compilation of palaeocurrent data across this transition is required. acknowledgements this work was undertaken during the wollaston forland regional mapping project, coordinated jointly by the geological survey of denmark and greenland (geus) and the greenland ministry of industry, energy, research and labour (mierl). the authors would like to thank the editor, mette olivarius, and tiffany playter and mihai emilian popa for their helpful reviews. additional information funding statement this project was jointly funded by geus and mierl. competing interests the authors declare no competing interests. author contributions sda: field data collection, writing and conceptualisation. hv: field data collection, writing and manuscript editing. jab-k: coal & bitumen analysis and writing. additional files there are no supplementary files with this manuscript. references ahokas, j.m., nystuen, j.p. & martinius, a.w. 2014: depositional dynamics and sequence development in a tidally influenced marginal marine basin: early jurassic neill klinter group, jameson land basin, east greenland. in: stevens, t. 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(2021). ba: bastians dal formation. mu: muslingebjerg formation. pe: pelion formation. fig. 2 location and geology map. (a) overview of the region, based on koch and haller (1971). (b) detailed geological map of the study region, modified from alsgaard et al. (2003, based on koch & haller (1971)) and using data collected during field mapping that formed part of this study. u: upper. m: middle. l: lower. fig. 3 photogrammetric overview of (a) the northern (bastians dal north) and (b) the southern (bastians dal south) areas of the study region. the locations of key observations and logged sections are indicated, along with the stratigraphic boundaries identified during this study. the numbering (i–ix) of logged sections in a relates to those illustrated in figs 4 and 7. fig. 4 summary log through the bastians dal formation (a) and detailed logs of the well exposed sections (b). the general position of the composite section is indicated in fig. 2. positions of the individual detailed logs, numbered i–viii, are provided in fig. 3a. relative positions as indicated on the summary log were calculated using photogrammetry. samples collected from these sections are numbered to the right of the logs. fig. 5 photographs illustrating the facies of the bastians dal formation. (a) large-scale cross-bedded conglomerates displaying low-angle truncation surfaces and bar-top relief (c. 69 m, fig. 4). (b) trough cross-bedded conglomerates with sharply erosive bases mantled with pebble (c. 115 m, fig. 4). (c) cross-bedded sandstones. clear alternations of coarser-and finer-grained foresets can be observed, as well as voids, where coalified wood was once present (c. 103.5 m, fig. 4) . fig. 6 compiled trough cross-bedding, planar cross-bedding and imbrication palaeocurrent data from the bastians dal formation (n = 18). data were collected from thorughout the outcrop area. fig. 7 sedimentological log and accompanying photograph of the thickest coal found in the muslingebjerg formation (ix in fig. 3a). this exposure is found in northern bastians dal. much thinner than its southern counterpart, the associated facies have a greater affinity to the underlying bastians dal formation. location in fig. 2. samples collected from this section are numbered to the right of the log. fig. 8 sedimentological log through the muslingebjerg formation in southern bastians dal. here, the succession is greatly expanded by the occurrence of shallow-marine sandstones. location and position of the log in figs 2 and 3b. samples collected from these sections are numbered to the right of the logs. fig. 9 photographs illustrating the facies of the muslingebjerg formation. (a) low-angle cross-bedded sandstones (c. 18 m, fig. 8). cross-bedding truncated by a horizontal erosion surface towards the top of the section. (b) bivalve moulds and bioturbation within the low-angle cross-bedded sandstones. fig. 10 rock-eval-type screening pyrolysis data for samples from bastians dal. for comparison, data from the present study are plotted together with data on coal samples from the same location, published by petersen et al. (2002). (a) standard-plot of tmax vs. hydrogen index (hi). samples from the present study all fall within the range of previously published samples. sample 541232 shows somewhat higher petroleum potential than samples 541231 and 541235. (b) standard-plot of toc vs. s2. samples from the present study all fall within the range of previously published samples. (c) linear plot of toc vs. s2. following the method of dahl et al. (2004), a regression line has been constructed for the lower, more prolific seam, suggesting high levels of inert carbon, but also high petroleum potential of the live kerogen fraction (hilive). sample 541232 falls on the regression line. (d) standard plot of tmax vs. production index (pi). only samples from the present study are shown. sandstone sample 541229 appears lightly stained, whereas 541228 does not. fig. 11 photomicrographs of coal sample 541232. width of field approximately 300 μm. (a) huminite particle (grey) enclosing lump of resinite (bright pinkish yellow). (b) field as a. fluorescence-inducing blue light. (c) huminite particle with minute fractures showing discolouration along the rims due to liquid bitumen. refraction anomalies akin to ‘newton’s rings’ are caused by bitumen from the fractures dissolving into the immersion oil used for microscopy. (d) same field as c. fluorescence-inducing blue light. note the presence of fluorescence of the petroleum-stained rims of the fractures and blooming phenomena at refraction anomalies. fig. 12 photomicrographs of sandstone sample 541229. width of field approximately 300 μm. (a and c) rounded quartz grains, pore surfaces lined with patchy, thin coatings of bitumen and carbonate cement. (b) same field as a. fluorescence-inducing blue light. (d) same field as c. fluorescence-inducing fig. 13 synthesis of sedimentological observations and stratigraphic geometries recorded in the bastians dal region. (a) infilling valley that existed during the deposition of the bastians dal formation. a gravel-bed fluvial system dominated deposition with finer-grained overbank facies becoming more common over time and including minor coal developments. the sw alignment of the valley may have been controlled by incipient faulting, which later also controlled the distribution of facies and stratal geometries within the muslingebjerg formation. palaeoflow of the fluvial system is indicated by the arrow. (b) interaction of environments recorded during the deposition of the muslingebjerg formation. rising relative sea level and reduced topography in the hinterland led to widespread coal formation (dark grey). these were overlain by shallow-marine shoreface and tidal (arrowed circulation: dashed line) sandstones as sea level continued to rise. with continued regional transgression, the muslingebjerg formation gave way to the shallow-marine sandstones of the pelion formation (not shown). palaeoflow of the fluvial system is indicated by the thin arrow (continuous line). as is clear from the stratal geometries described in this study, subsidence during this period was fault controlled. legend for sedimentary structures in figs 4 and 8. table table 1 results of coal and bitumen analysis. geological survey of denmark and greenland bulletin 14, 78 pp. geological survey of denmark and greenland bulletin 14 · 2007 quaternary glaciation history and glaciology of jakobshavn isbræ and the disko bugt region, west greenland: a review anker weidick and ole bennike geological survey of denmark and greenland ministry of climate and energy bulletin 14: gsb191-indhold 04/12/07 14:36 side 1 geological survey of denmark and greenland bulletin 14 keywords jakobshavn isbræ, disko bugt, greenland, quaternary, holocene, glaciology, ice streams, h.j. rink. cover mosaic of satellite images showing the greenland ice sheet to the east (right), jakobshavn isbræ, the icefjord kangia and the eastern part of disko bugt. the position of the jakobshavn isbræ ice front is from 27 june 2004; the ice front has receded dramatically since 2001 (see figs 13, 45) although the rate of recession has decreased in the last few years. the image is based on landsat and aster images. landsat data are from the landsat-7 satellite. the aster satellite data are distributed by the land processes distribution active archive center (lp daac), located at the u.s. geological survey center for earth resources observation and science (http://lpdaac.usgs.gov). frontispiece: facing page reproduction of part of h.j. rink’s map of the disko bugt region, published in 1853. the southernmost ice stream is jakobshavn isbræ, which drains into the icefjord kangia; the width of the map illustrated corresponds to c. 290 km. chief editor of this series: adam a. garde scientific editor of this volume: jon r. ineson editorial secretaries: jane holst and esben w. glendal referees: ole humlum (norway) and carl benson (usa) illustrations: stefan sølberg and henrik klinge pedersen digital photographic work: benny m. schark layout and graphic production: annabeth andersen printers: schultz grafisk, albertslund, denmark manuscript received: 11 may 2006 final version approved: 12 january 2007 printed: 17 december 2007 issn 1604-8156 isbn 978-87-7871-207-3 geological survey of denmark and greenland bulletin the series geological survey of denmark and greenland bulletin replaces geology of denmark survey bulletin and geology of greenland survey bulletin. citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 14, 78 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk or at www.geus.dk/publications/bull © de nationale geologiske undersøgelser for danmark og grønland (geus), 2007 for the full text of the geus copyright clause, please refer to www.geus.dk/publications/bull bulletin 14: gsb191-indhold 04/12/07 14:36 side 2 bulletin 14: gsb191-indhold 04/12/07 14:36 side 3 4 contents abstract . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5 introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7 jakobshavn isbræ and disko bugt . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7 setting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7 areas and volumes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8 mass balance of the inland ice . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9 gross features of the coastland and major drainage of the ice sheet . . . . . . . . . . . . . . . . . . . . . 11 present climate . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 12 history and exploration . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 14 archaeology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 14 discovery, rediscovery, early mapping and descriptions up to c. 1845 . . . . . . . . . . . . . . . . . . . 14 hinrich johannes rink (1819–1893) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 18 observations and mapping of the ice margin around disko bugt . . . . . . . . . . . . . . . . . . . . . . 21 large-scale glaciological projects after world war ii . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 22 deep cores from the ice sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 24 hydropower and climatic change . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 24 geology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 26 bedrock geology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 26 history of glaciations and interglacials . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 28 the last interglacial (sangamonian/eemian) and the last ice age (wisconsinan/weichselian) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 30 the collapse of the ice cover in disko bugt . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 35 the ‘fjord stage’ and the attainment of the present ice-margin position . . . . . . . . . . . . . . . . . 37 the ice-sheet margin after the holocene thermal maximum . . . . . . . . . . . . . . . . . . . . . . . . . . 42 relative sea-level changes around disko bugt . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 46 glaciology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 50 subsurface of the ice-sheet margin . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 50 present ice-margin surface . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 51 drainage and thermal conditions of the ice margin . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 54 movement of the ice margin . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57 notes on individual outlets . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 59 summary and outlook . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 65 acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 66 references . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 67 appendix 1: index to greenland place names . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 76 appendix 2: radiocarbon analyses . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 78 bulletin 14: gsb191-indhold 04/12/07 14:36 side 4 5 abstract weidick, a. & bennike, o. 2007: quaternary glaciation history and glaciology of jakobshavn isbræ and the disko bugt region, west greenland: a review. geological survey of denmark and greenland bulletin 14, xx pp. the disko bugt region in central west greenland is characterised by permanent ice streams, of which jakobshavn isbræ is by far the most important. the first thorough studies on the glaciology of the region were conducted over 150 years ago by h.j. rink, who introduced the terms ‘ice streams’ and ‘inland ice’. rink’s work inspired new field work, which has continued to the present, and the long series of observations are unique for an arctic region. cooling during the cenozoic led to ice-sheet growth in greenland. a number of interglacial occurrences have been reported from the disko bugt region, and during the penultimate glacial stage, the greenland ice-sheet margin extended to the shelf break. during the last glacial maximum, the ice margin probably extended only to the inner part of the banks on the continental shelf, and large floating glaciers may have been present at this time. during the younger dryas cold period, the ice margin may have been located at a marked basalt escarpment west of disko bugt. disko bugt was deglaciated rapidly in the early holocene, around 10 500 – 10 000 years before present (10.5–10 ka b.p.), but when the ice margin reached the eastern shore of the bay, recession paused, and major moraine systems were formed. with renewed recession, the present ice-margin position was attained around 8–6 ka b.p., and by c. 5 ka b.p. the ice margin was located east of its present position. the subsequent neoglacial readvance generally reached a maximum during the little ice age, around ad 1850. this was followed by recession that has continued to the present day. the relative sea-level history shows a rapid sea-level fall in the early holocene, and a slow rise in the late holocene. this development mainly reflects a direct isostatic response to the ice-margin history. jakobshavn isbræ is the main outlet from the greenland ice sheet. it drains c. 6.5% of the present inland ice, and produces c. 35–50 km3 of icebergs per year, corresponding to more than 10% of the total output of icebergs from the inland ice. the velocity of the central part of the ice stream at the front has been around 7 km/year since records began, but has nearly doubled in recent years. other calf-ice producing glacier outlets in disko bugt produce c. 18 km3 per year. the large calf-ice production of jakobshavn isbræ may have been initiated at about 8 ka b.p. when the glacier front receded from the iceberg bank (isfjeldsbanken) near ilulissat. ice streams in inner and outer egedesminde dyb may have been active during the early holocene and during the last glacial maximum. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: awe@geus.dk; obe@geus.dk bulletin 14: gsb191-indhold 04/12/07 14:36 side 5 6 nanortalik ammassalik kangerlussuaq i n l a n d i c e svalbard iceland swiss station gunnbjørn fjeld camp century arctic ocean fram strait jøkelbugten kap farvel dye 3 labrador sea svartenhuk halvø baffin island ellesmere island canada peary land kane basin thule air base (pituffik) harald moltke bræ gade gletscher sanddalen storstrømmen northgrip upernavik isstrøm rink isbræ upernavik melville bugt arfersiorfik davis strait kangerluarsunnguaq paamiut eqalorutsit killiit sermiat qaqortoq narsarsuaq narsap sermia sisimiut nuuk summit uummanaq fjord kangerlussuaq gletscher scoresby sund helheimgletscher kangerlussuaq airport ussuit kangerlussuaq baffin bay pattorfik qilertinnguit disko ilulissat ice survey stations, core sites towns 500 km fig. 1. map of greenland showing the localities mentioned in the text. bulletin 14: gsb191-indhold 04/12/07 14:36 side 6 7 jakobshavn isbræ (sermeq kujalleq) in disko bugt, west greenland, has been recognised as the king of greenland glaciers amongst scientists and travellers in the arctic for many decades. it is generally assumed that none of the other fast-moving outlets of the inland ice produce comparable quantities of icebergs. in december 2000, the greenland home rule authority decided to nominate the icefjord in front of jakobshavn isbræ together with the surrounding areas for inclusion in the world heritage list of unesco (united nations educational, scientific and cultural organisation). the nomination report was submitted in 2003 and ‘ilulissat icefjord’ was included on the world heritage list at the annual meeting of the world heritage committee in june 2004. this volume presents a comprehensive description of the region around jakobshavn isbræ, including the ‘ilulissat icefjord’ world heritage site, and emphasises the importance of the region for glaciological investigations in greenland. jakobshavn isbræ and disko bugt preparation of the ilulissat icefjord nomination report (mikkelsen & ingerslev 2002) involved perusal of the large number of scientific papers and descriptions of jakobshavn isbræ, its ice production and the glaciological and quaternary history of the region. however, the nomination document was a technical report published in a limited number, and following inclusion in the world heritage list a profusely illustrated book designed for a wider audience was produced and published in separate danish, english and greenlandic editions (bennike et al. 2004). the present volume documents the scientific background for the description and conclusions provided by bennike et al. (2004) in the above book. the historical introduction is followed by sections that focus on the geological history, the special peculiarities of the ice cover, and the present-day status of the glacier. the increasing number of recent publications, and the wide spectrum of scientific investigations of the glacier and its environment, have necessitated an updating of the descriptive section. the opportunity is taken here to discuss and present wider conclusions on the geological history of the ice sheet and its surroundings. in this volume, greenlandic place names are used according to the spelling that was introduced in 1973. however, the pre-1973 spelling is retained for established stratigraphic and other terms introduced prior to this year. the danish name of the main glacier in disko bugt used in most published descriptions is ‘jakobshavn isbræ’ (older version: jakobshavns isbræ), derived from the former name for the town on the north side of the fjord (jakobshavn, now ilulissat). the authorised greenlandic name for the gla cier is sermeq kujalleq (‘the southern glacier’), but this place name is also used for several other large outlet gla ciers that drain from the inland ice into disko bugt (qeqertarsuup tunua) and uummannaq fjord (uumman nap kangerlua). to avoid confusion with other glaciers in the area with the same name, the former name ‘jakobshavn isbræ’ is retained here for the glacier, a usage that accords with most published descriptions. the formal present-day name of the icefjord in front of the glacier is kangia (= ‘its eastern part’, i.e. east of the town of ilulissat), but other names are also used such as ‘ilulissat isfjord’, ‘jakobshavn isfjord’ and ‘jakobshavn icefjord’. the world heritage list uses the name ‘ilulissat icefjord’ for the entire world heritage site, which includes land areas and parts of the inland ice adjacent to the icefjord. a few other danish place names are used in order to avoid confusion or for historical reasons. greenland place names used are listed in the locality index (appendix 1) at the end of this treatise; the locations of place names appear on figs 1, 2. setting the disko bugt region including jakobshavn isbræ is situ ated in the central part of west greenland, with the position of the present front of jakobshavn isbræ located at c. 69º10´n, 50ºw. as the glacier is fed by an extensive sector of the ice sheet, a summary of the general morphologi cal and glaciological features of greenland (fig. 3) is provided as a background for the following description. compre hensive descriptions of the present physiography of greenland and its relation to the complex geological history are provided by escher & watt (1976), funder (1989) and henriksen et al. (2000). a comprehensive account of the dynamic and climatic history of the inland ice is given by reeh (1989). the importance of the major ice stream of jakobshavn isbræ can best be illustrated by the size of the sector of the inland ice that feeds it (fig. 3). this was estimated at introduction bulletin 14: gsb191-indhold 04/12/07 14:36 side 7 8 between 3.7 and 5.8% of the inland ice, corresponding to an area of 63 000 – 99 000 km2, by bindschadler (1984). more recent estimates have increased this figure to 6.5% of the ice sheet, i.e. 110 000 km2 (echelmeyer et al. 1991). this extensive catchment area accounts for the greater part of the actual ice flow to the disko bugt region, based on zwally & giovinetto (2001). areas and volumes the total area of greenland is c. 2.2 million km2 of which the ice sheet (the inland ice) constitutes c. 1.7 million km2 (weng 1995). this latter figure also includes some minor marginal ice caps that, while contiguous with the inland ice proper, have their own ice dynamics. these ice caps are situated on highlands (especially in east greenland) and usually have a thickness of a few hundred metres (weidick & morris 1998); their combined area only amounts to a few per cent of the total area of the inland ice. the inland ice is an approximately lens-shaped body, with a maximum thickness of 3.4 km, and with the highest elevation at summit of 3238 m a.s.l. (fig. 3). the inland ice rests in a bowl-shaped depression (fig. 4), which in the central parts is below present sea level due to the depression of the earth’s 70°n 69°n 54°w 25 km 52°w 50°w jakobshavn isbræ kangilerngata sermia tors ukattak sikuiuitsoq nordenskiöld gletsc her eqip sermia qapiarfiit sermeq kujalleq sermeq avannarleq qarajaq isfjord sermeq kujalleq disko bugt disko qeqertarsuaq (godhavn) disko fjord tuapaat innaarsuit aasiaat tininnilik saqqarliup sermia tasiusaq narsarsuaq tissarissoq oqaatsut qajaa vaskebugt kangia se rm eq av an na rle q arveprinsen ejland paakitsoq alanngorliup sermia nuuk eqaluit kangersuneq ilimanaq akulliit isfjeldsbanken sermermiut orpissooq ilulissat qasigiannguit qeqertarsuatsiaq naternaq hareøen vaigat nuussuaq nordfjord fig. 2. map of the disko bugt region showing the localities named in the text. the stippled line indicates the position of isfjeldsbanken – the shoal at the mouth of kangia. bulletin 14: gsb191-indhold 04/12/07 14:36 side 8 9 crust caused by the load of the ice. estimates of the volume of the inland ice vary from c. 2.6 million km3 (holtzscherer & bauer 1954) to 2.9 million km3 (bamber et al. 2001; layberry & bamber 2001), corresponding to c. 7% of the world’s fresh water (reeh 1989). mass balance of the inland ice snow accumulation in the central parts of the ice sheet and loss in the marginal parts govern the mass balance of the inland ice. the accumulation is estimated to be 500–600 km3 ice per year, which until 2000 was assumed to approxi mately match the loss. about half of the loss was ascribed to melting, and the other half to calving. bottom melting of floating glaciers may reduce the calf-ice production of north and north-east greenland outlets (reeh 1989, 1994, 1999). although calving glaciers are widespread along the coasts of greenland, the main annual loss of calf ice from the inland ice is concentrated at a rather small number of outlets along its c. 6000 km long perimeter. many of the important calving glaciers are found along the west coast of greenland, with approximately 84 km3 of the calf-ice production originating from five outlets (table 1, fig. 5). four of these occur along a 300 km stretch of coast in the disko bugt – uummannaq fjord region. the calf-ice production of jakobshavn isbræ is of particular importance for the mass balance of the ice sheet (fig. 5). the production has generally been estimated to be about 35 km3 ice per year, corresponding to more than 10% of the estimated total output of icebergs from the inland ice, but more recent estimates are higher, around 50 km3 per year in 2003, according to joughin et al. (2004). jakobshavn isbræ is considered the most active glacier in greenland by legarsky & huang (2006). the reason for the large ice production is that a major part of the inland ice drains towards the central part of west greenland, especially the relatively low uplands in the interior of disko bugt (fig. 3). with regard to the dynamics of glaciers, determinations of movement and calf-ice production over long time spans are rare. for jakobshavn isbræ and sermeq avannarleq in torsukattak velocity measurements go back to 1875. it appears from scattered velocity records for the disko bugt jakobshavn isbræ gunnbjørn fjeld dye 3 camp century summit northgrip ice divide deep ice core n 250 km fig. 3. map of greenland showing the location of jakobshavn isbræ and the deep cores on the inland ice. the approximate ice drainage area to jakobshavn isbræ is shown with the solid red line, and the ice drainage area to the entire disko bugt region with the dashed line; these are based on the flow-line map of zwally & giovinetto (2001). the dotted red line shows the trend of the ice divide. summit (3238 m a.s.l.) is the highest point on the ice sheet, and gunnbjørn fjeld (3693 m a.s.l.) is the highest mountain in greenland. original map base by simon ekholm, reproduced with the permission of kort & matrikelstyrelsen (kms) [national survey and cadastre], copenhagen. glacier latitude production velocity (km3/year) (km/year) jakobshavn isbræ (sermeq kujalleq) 69°11´n c. 35 5–7 sermeq kujalleq in torsukattak 70°00´n 8–10 2.6–3.5 sermeq kujalleq (store gletscher) 70°20´n 14–18 4.2–4.9 rink isbræ (kangilliup sermia) 71°45´n 11–17 3.7–4.5 gade gletscher 76°20´n c. 10 sources: bauer et al. (1968a); carbonell & bauer (1968); weidick (1995). table 1. calf-ice production from the five largest outlets in west greenland bulletin 14: gsb191-indhold 04/12/07 14:36 side 9 10 region from the last part of the 19th and the 20th century that the glaciers have maintained a rather permanent rate of flow. in other parts of greenland, outlets from the ice sheet show pulsating or surging behaviour, such as harald moltke bræ (ullip sermia) in north-west greenland (mock 1966), storstrømmen in north-east greenland (bøggild et al. 1994) and eqalorutsit killiit sermiat in south greenland (weidick 1984). more recent examples are provided by rignot & kanagaratnam (2006). over the past decades, aircrafts and satellites are increasingly being used to monitor the inland ice. this has led to more detailed observations on changes of velocity, calf-ice production, ice elevation and frontal positions. this development coincided with dramatic changes in the marginal parts of the ice sheet. marked thinning and recession have been reported for many outlets in greenland (rignot & kanagaratnam 2006). the velocity of many glaciers has increased, and the ice-sheet mass deficit changed from 90 to 220 km3 per year between 1996 and 2005. in east greenland, kangerlussuaq gletscher (fig. 1) accelerated 210% from 2000 to 2005, and its front receded 10 km. with its velocity of 13–14 km/year it is now the fastest gla cier in greenland. helheimgletscher farther south accelerated 60% and receded 5 km. in 2001, the velocity of this glacier was measured to be c. 8 km/year (thomas et al. 2001). in west greenland, narsap sermia accelerated by 150% from 2000 to 2005 while jakobshavn isbræ accelerated by 95% and receded c. 10 km. the velocity of jakobshavn isbræ was 12.6 km/year in 2003 (joughin et al. 2004). the marked flow-velocity increase is considered to be related to global warming, which leads to increased melting and sliding in the marginal parts of the ice sheet (krabill et al. 2000, 2004). a thinning of around 2 m per year is reported for marginal parts of the inland ice, which is scarcely matched by a snow accumulation increase of 5–6 cm per year in the central parts of the ice sheet (alley et al. 2005; johannesen et al. 2005; dowdeswell 2006; rignot & kanagaratnam 2006). however, whereas jakobshavn isbræ has maintained the high frontal velocity for several years, the velocity increase of other glaciers (kangerlussuaq gletscher and helheimgletscher in eastern greenland) seems to have been a short-lived event as the velocity and discharge have decreased since 2006 (howat et al. 2007; truffer & fahnestock 2007). the net loss of ice from the greenland ice cover plays an important role in global sea-level rise, and therefore more detailed investigations of the causes for the marked changes in greenland are required to assess and model ongoing and future changes. the recently observed changes may lead to a new stable situation, but if the changes continue they may eventually lead to the disappearance of the inland ice (alley et al. 2005). however, we note that the inland ice did not disappear during the eemian, even though temperatures were around 5°c higher than at the present. n 250 km elevation of bedrock (m) 900–3300 800–900 700–800 600–700 500–600 400–500 300–400 200–300 100–200 0–100 –100–0 –200 to –100 –300 to –200 –400 to –300 –1000 to –400 fig. 4. bedrock topography below the inland ice, from bamber et al. (2001) and layberry & bamber (2001); data provided by the national snow and ice data center daac, university of colorado, boulder, usa. elevations are not corrected for the present load of the glacier ice (cf. fig. 17). the map shows the ‘channel’ connecting the subglacial central basin to jakobshavn isbræ and the disko bugt region. bulletin 14: gsb191-indhold 04/12/07 14:36 side 10 11 gross features of the coastland and major drainage of the ice sheet at the present day, the inland ice is separated from the sea by a coastal strip of more or less ice-free land, which is up to 300 km wide. the outlets from the inland ice to the sea are, for the most part, restricted to fjords that cut through the coastal strips of ice-free land. exceptions are found in north-west greenland, in melville bugt (qimusseriarsuaq) and in the kane basin, where large parts of the ice sheet margin reach the sea. the high alpine mountains of eastern greenland, with peaks up to 3693 m (gunnbjørn fjeld at 68º55´n, 29º53´w, fig. 3) form a topographic barrier that is in contrast to the widespread hilly uplands of western greenland. the general topographic fall from this high 40 km3/year 30 20 10 0 gade gletscher (9.8 km3/year) sermeq (upernavik isstrøm) salliarutsip sermia (ingia isbræ) umiammakku sermia (umiamako isbræ) kangilliup sermia (rink isbræ) kangilleq sermilik sermeq silarleq kangerlussuup sermersua kangerluarsuup sermia perlerfiup sermia sermeq avannarleq (lille gletscher) sermeq kujalleq (store gletscher) sermeq kujalleq in torsukattak sermeq avannarleq in torsukattak kangilerngata sermia eqip sermia sermeq avannarleq at ilulissat alanngorliup sermia saqqarliup sermia akuliarutsip sermersua (nordenskiöld gletscher) usulluup sermia narsap sermia akullersuup sermia nakkaasorsuaq avannarleq bræ uukkaasorsuaq sermiligaarsuk bræ sermeq in arsuk fjord sermilik bræ qalerallit sermia eqalorutsit killiit sermiat qooqqup sermia qajuuttap sermia (eqalorutsit kangilliit sermiat) kangiata nunaata sermia sermeq kujalleq (jakobshavn isbræ) fig. 5. estimated maximum (dark blue) and minimum (pale blue) calf-ice production from glaciers along the west coast of greenland. jakobshavn isbræ (sermeq kujalleq) is clearly in a class of its own (from weidick et al. 1992). the data are based on bauer et al. (1968a) and carbonell & bauer (1968), and hence do not consider the dramatic change in calf-ice production after c. a.d. 2000. bulletin 14: gsb191-indhold 04/12/07 14:36 side 11 12 mountainous barrier is to the west, and thus the main ice drainage is towards the west. offshore greenland, continental shelves and slopes of variable width border the deep sea. the width of the shelf varies from 50 to 300 km, with the greatest extent off jøkelbugten in north-east greenland, off kangerlussuaq in south-east greenland, and off disko bugt in west greenland. the water depth over the shelf is mainly between 100 and 300 m, except off south-east greenland where large shelfs areas are found at 300–400 m below sea level. the total area of the continental shelf is estimated to be c. 825 000 km2 (henriksen et al. 2000). sinuous troughs, 600–1000 m deep, cut through the shelf at intervals and extend from fjord systems or depressions to the continental slope. their origin is connected with drowned valleys of neogene age (pelletier 1964) that were over-deepened and transformed by glacial erosion during the ice ages. quaternary marine sediments and till cover large areas of the shelf. ridges and areas of coarse sediments on the banks are interpreted as marginal deposits that were formed at advanced positions of the ice-sheet margin during the ice ages. present climate a high pressure system usually prevails over the greenland inland ice at the present day. the high landmass of greenland tends to split eastward-moving low pressure systems approaching from the south-west into two separate depressions, one travelling north along west greenland and one travelling up along the east coast of greenland. in addition, a number of depressions also approach west greenland from the hudson bay region; most precipitation in west greenland is related to the passage of low pressure systems (hansen 1999). the climate along the outer coast in eastern and southern greenland is affected by the east greenland current that transports large amounts of sea ice from the arctic ocean down along the coast of east greenland, around kap farvel, and northwards up to the paamiut area at c. 62ºn on the west coast. due to the relatively warm current washing certain stretches of west greenland, parts of the coastal region are open for ship traffic, even in winter. the disko bugt region and areas to the north are usually covered by sea ice during the winter. in the summer, western greenland up as far as the thule area is nearly ice-free, with the exception of calf ice from the inland ice outlets (buch 2002). precipitation in greenland shows a marked general decrease from south to north. more than 2500 mm/year is recorded in southern greenland, decreasing to less than 150 mm/year in the interior of peary land in northern greenland (reeh 1989). in all parts of greenland, a decrease in precipitation is apparent from the outer coast to areas near the ice-sheet margin. the latter areas are characterised by a continental type of climate; mean annual precipitation at kangerlussuaq airport in west greenland at 67ºn is around 150 mm, and at ilulissat at 69º13´n it is around 270 mm (cappelen et al. 2001). on the western slope of the inland ice at around 70º46´n, observations indicate 1880 1900 1920 1940 1960 year (a.d.) 1980 2000 ilulissat te m pe ra tu re ( °c ) nuuk pituffik narsarsuaq upernavik 2 –2 –4 –6 –8 –10 –12 –14 0 fig. 6. annual mean temperatures between 1873 and 2004 for five stations in west greenland, including ilulissat (redrawn from cappelen 2005). the smooth curves were created using a gauss filter; locations of the stations shown on fig. 1. bulletin 14: gsb191-indhold 04/12/07 14:36 side 12 13 a mean annual precipitation of c. 600 mm that is thought to be related to the relatively easy access of humid air masses passing through disko bugt (reeh 1989); this results in lowering of the glaciation level in this region (humlum 1985, 1986). the average accumulation for the entire ice sheet is around 31 cm water equivalent annually (ohmura & reeh 1991), with a general south to north decrease. in west greenland, the mean annual air temperature decreases from +0.6ºc at qaqortoq (61ºn) to –3.9ºc at ilulissat (69ºn) and –11.1ºc at thule air base at 77ºn (cappelen et al. 2001; box 2002). these values are based on data from meteorological stations situated close to sea level. historical data from these and other meteorological stations show a rise in temperature between c. 1880 and c. 1950, followed by cooling of 1–2ºc between c. 1950 and c. 1990 and a subsequent marked warming (fig. 6). data from a network of automatic weather stations indicate that the mean annual air temperature for the central parts of the ice sheet was around c. 2ºc higher during the time period 1995–1999, as compared to the period 1951–1960. at an elevation of 1000–2000 m, the temperature increase was only c. 1ºc (steffen & box 2001). on the ice sheet itself, the mean annual temperature at the surface varies between 0ºc and –32ºc, according to elevation and geographical location. the temperature increases with depth in the ice body and can reach the pressure-melting point (–2.6ºc for 3000 m thick ice) due to geothermal heat and heating caused by ice deformation. it is generally believed that most glacial erosion takes place near the margins of the ice sheet. however, cold bed conditions are found at high, elevated areas with thin and almost stagnant ice. in such areas, the ice cover will protect the subsurface, rather than erode it (reeh 1989). bulletin 14: gsb191-indhold 04/12/07 14:36 side 13 14 archaeology in the disko bugt region, the three waves of palaeo-eskimo and neo-eskimo cultures in west greenland (fig. 7) are richly represented by numerous archaeological sites (larsen & meldgaard 1958; j. meldgaard 1983; m. meldgaard 2004; gulløv et al. 2004). remains from palaeo-eskimo and neo-eskimo cultures can be found as far inland as qajaa, which was situated only a few kilometres from the front of jakobshavn isbræ during its maximum extent at around 1850. the oldest culture, the saqqaq culture, is dated to the period from c. 2500 b.c. to about 800 b.c. at sermermiut, the deposits of this culture are separated from those of the subsequent dorset culture by a sterile layer. during the time of the greenlandic dorset culture (also referred to as the early dorset culture, c. 800 b.c. to 0 b.c.), a smaller and more scattered population lived in the region around disko bugt. at about a.d. 1100, the third (thule) culture arrived in greenland and settled in the disko bugt region about 100 years later (gulløv et al. 2004). the uninhabited periods between the three cultures have been related to climatic deterioration, but other factors such as overexploitation of the natural resources may also have played a role. during some periods, the disko bugt region was densely populated, judging from the large number of settle ment sites – at least compared to other regions in the arctic. the well-known former settlement, the sermermiut prehistoric ‘town’ situated close to present-day ilulissat, was abandoned in the middle of the 1800s. the establishment of the trade centre at jakobshavn (ilulissat) by jakob severinsen in 1741 was probably the main factor behind the abandonment of sermermiut. although there was certainly some contact between the thule neo-eskimo settlers that migrated from the north and the norse people that came from the south, the surviving icelandic sagas provide no identifiable description of the region around ilulissat. discovery, rediscovery, early mapping and descriptions up to c. 1845 the first description of the ice cover of greenland comes from the norse people, who settled in southern west green land after the arrival of erik the red in a.d. 986 (fig. 7; gad 1967, p. 43). in ‘kongespejlet’ (‘the king’s mirror’, here cited from the english translation by larson 1917, p. 143–144), a description written about 1260 records: “in reply to your question whether the land thaws or remains icebound like the sea, i can state definitely that only a small part of the land thaws out, while all the rest remains under the ice. but nobody knows whether the land is large or small, because all the mountain ranges and all the valleys are covered with ice, and no opening has been found anywhere”. although this description concerns southern greenland, it is generally valid for the whole of greenland. the norse settlements were abandoned in the 1400s (gad 1967; arneborg 1996). the disappearance of the norse population is often linked with the onset of the little ice age although other factors, such as exhaustion of natural resources, rising sea level, the political situation in history and exploration 2500 b.c. 2000 1500 1000 norse (a.d. 985–1450) colonisation 500 500 0 1000 1500 a.d. 2000 ye ar independence i (2500–1900 b.c.) thule (a.d. 1100 – present) saqqaq (2500–800 b.c.) greenlandic dorset (800 b.c. – 0) late dorset (a.d. 700–1300) fig. 7. chronology of the different archaeological cultures that have colonised greenland according to gulløv et al. (2004). modified from bennike et al. (2004). bulletin 14: gsb191-indhold 04/12/07 14:36 side 14 15 fig. 8. poul egede’s map of greenland from 1788, showing the fictitious strait extending across greenland from disko bugt to the coast of east greenland (from nordenskiöld 1886, p. 212). the size of the original map is 294 × 379 mm. bulletin 14: gsb191-indhold 04/12/07 14:36 side 15 16 scandinavia, the black death (plague), or attacks by eskimos or biscay pirates, may also have played a role (gad 1967). apart from sporadic observations by english and dutch whalers and explorers, there are no descriptions of the physi ography of greenland or its ice cover during the 1500s and 1600s. for disko bugt in particular, the survival of numerous place names of dutch origin (rodebay, claushavn, vaigat etc.) is linked to dutch whaling activity in the 1600s and the beginning of the 1700s. an early description of the icefjord was published by the dutch whaler, feykes haan, in a navigation guide: “half a mile north of sant-bay is a fjord that is always full of ice, with frightfully tall icebergs, but from where they come is unknown. this fjord is called ys-fioert [icefjord].” (haan 1719, quoted from bobé 1916, p. 47; authors’ translation). short as it is, this description reveals that the conditions of the iceberg bank and the icefjord in the early 1700s were much the same as can be seen today. up to the beginning of the 1700s, knowledge of the west greenland coastal region was poor. in 1721, however, the priest hans egede settled in greenland near godthåb (now nuuk, at c. 64º10´n), and in subsequent years permanent trading stations and missions spread rapidly. by the end of the 1700s, a network of danish settlements covered west greenland from nanortalik in the south (at 60ºn, established 1797) to upernavik in the north (at 72ºn, established 1772). in the disko bugt region, qasigiannguit (christianshåb) was established in 1734, ilulissat (jakobs havn) in 1741, aasiaat (egedesminde) in 1763 and qeqer tarsuaq (godhavn) in 1773. with colonisation came the first attempts to undertake a systematic mapping of greenland (in particular west greenland) and also an increasing number of local descriptions were made. the development of early mapping is described and illustrated by dupont (2000). poul egede’s map from 1788 illustrates the status of mapping at that time (fig. 8). apart from the lack of detail, it should be noted that a channel is depicted connecting disko bugt in west greenland with kangerlussuaq in east greenland. this socalled ‘frobisher strait’ is an error copied from older maps. egede commented on the map: “it is said that the strong current that flows continuously from the ice-dome comes from ollum længri fjord” (i.e. from east greenland; authors’ translation). egede also stated in another comment to the map: “the entire land is concealed under ice and snow, from staten huk to the extreme north” (authors’ translation). ‘staten huk’ was the dutch whalers’ name for kap farvel at the southern tip of greenland. the way this is expressed perhaps indicates some doubt about the alleged channel through greenland. egede’s comments about the icefjord, with respect to the strong currents from the glacier may be based on observation, as may his depiction of the icefjord with a length similar to the present. historical descriptions of the conditions in kangia and its surroundings were collated by larsen & meldgaard (1958), supplemented by georgi (1960a, b). both stressed the records that the icefjord was more ice-free in the early 1700s than in subsequent times. this conclusion is essentially based on a letter from the manager of jakobshavn, hans rosing, from 1831. it states in translation: “an old woman still living here [at jakobshavn] knew, when she was young, another old woman who told her that, when she was a young girl, there were practically no icebergs in the fjord, but the water was so open that the dutchmen with their large vessels went into the fjord. up along the coast of the fjord the greenlanders lived in their tents, and in the winter there were houses of which ruins still can be seen at least one ‘miil’ [approx. 10 km] from the mouth” (larsen & meldgaard 1958, p. 24–25). historically these events can be related to a time just before 1740. this evidence can be combined with descriptions from a visit to the ice margin on 16 january 1747 by p.o. walløe (bobé 1927). walløe described the advancing ice margin in so much detail that there can be little doubt of the validity of his observations, which were probably made close to qajaa (fig. 9). qajaa was abandoned in the middle of the 1700s, presumably because of the advance of jakobshavn isbræ in the 18th and first half of the 19th century. rink (1875, p. 15) provides further confirmation: “…proof that jakobs havn-fjorden was earlier accessible further in, is given by the remains of an older dwelling site in a location that cantasiusaq nunatarsuaq qajaa sermermiut 5 km kangia ilimanaq (claushavn) ilulissat (jakobshavn) 1851 sikuiuitsoq 51°30´w 69°15´n fig. 9. map of kangia and the surrounding region, showing the location of former inuit settlements (indicated by the red dots) and the approximate position of the front of jakobshavn isbræ (sermeq kujalleq) in a.d. 1851 (white line). compiled by the national museum and archives of greenland; modified from bennike et al. (2004). bulletin 14: gsb191-indhold 04/12/07 14:36 side 16 17 a b c fig. 10. a: rink’s concept of the drainage of the interior of greenland by rivers (rink 1862). the outline of east greenland was then not well known, and the depiction of the area north of ilulissat in west greenland is also imprecise. b: rink’s map showing the extent of the inland ice, and the large ice streams that drain into disko bugt and uummannaq fjord (rink 1857). smaller glaciers in southern greenland are also indicated. c: segment of rink’s original map of the interior part of disko bugt showing the positions of the ice streams draining into kangia and torsukattak (rink 1853). bulletin 14: gsb191-indhold 04/12/07 14:36 side 17 18 not be reached now due to ice” (authors’ translation). together these observations lead to the conclusion that the glacier could have been in a rather retracted position in 1747, but perhaps in an initial advancing state. the gla cier may well have been just as productive then as it is now, as implied by haan’s description of icebergs on the iceberg bank in 1719. an initial advance in the beginning of the 1700s may only have led to a slight reduction in calf-ice production. unfortunately, the observations on the advancing ice margin around jakobshavn isbræ up to c. 1850, backed up by records of vegetated soil or ruins buried by the advan cing ice, generally lack information as to the exact location at which they were made. these observations from the 18th and the beginning of the 19th century, however, form valuable contributions to the growing interest in the changes of ice cover. the observations on the greenland ice cover from the 18th century were compared by cranz (1770) with the status of european and american glaciers, which were then also advancing. this provided impetus for more detailed mapping and descriptions of the greenland ice cover in the following century, when it was first appreciated that glaciers could be regarded as a kind of ‘climatoscope’, with their recessions and advances reflecting alternating warm and cold periods. in greenland, systematic observations and descriptions of the ice cover were initiated by hinrich johannes rink and recorded in a series of outstanding publications. hinrich johannes rink (1819–1893) originally educated in chemistry and physics (oldendow 1955), h.j. rink carried out geological investigations in parts of northern west greenland between 1848 and 1852. from 1853 to 1858 he was trade manager in qaqortoq (julianehåb) and nuuk (godthåb), and from 1858 to 1868 he was inspector of the royal greenland trade in south greenland. he travelled over large parts of west greenland. one of his ambitions was to produce a pioneer map of west greenland, incorporating his own mapping with that of older sea charts, the mapping of the missionary samuel kleinschmidt (1814–1886; see wilhjelm 2001, p. 144–152), and map sketches made by local hunters. he devoted much of his own mapping efforts to the interior, eastern ice-free fjords and land areas that were largely unknown at the time. during his travels he visited, described and mapped extensive areas of the icefjords and their glaciers. rink stayed in ilulissat (jakobshavn) over the winter of 1850–1851, and sailed to the inland ice margin at the fjord paakitsoq, north of ilulissat, in october 1850. he also travelled by dog sledge to the southern side of the front of jakobshavn isbræ in april 1851 (rink 1857; vol. 1) and visited the north side of the glacier in may the same year (von drygalski 1897, p. 129). his approximate determination of the frontal position of jakobshavn isbræ is shown in fig. 9, as depicted in later reviews. the original version of rink’s map is shown in fig. 10. rink’s observations of the frontal position of jakobshavn isbræ in 1850–51 were the first in a series of observations that have now extended over more than 150 years. the long series of observations of the frontal changes and the velocity of jakobshavn isbræ (from 1875), and the long series of continuous meteorological records at ilulissat (since 1873), are unique for an arctic area. rink was the first scientist to appreciate the immense extent and special form of the ‘ice plain’, covering the entire interior region of greenland. rink called it ‘indlandsisen’ [the inland ice] following a suggestion by the danish scientist japetus steenstrup. it was clear to rink that this large body of ice was quite different from the local glaciers hitherto described from other parts of the world. in europe a new idea was emerging at this time, namely that extensive ice sheets had covered large parts of northern europe in the past. rink’s demonstration of an extant, immense ice sheet in greenland was sensational news, and provided critical support for arguments that such an ice cover had once existed in europe. while it is true that earlier mapping and descriptions, such as poul egede’s map (fig. 8), had indicated the presence of extensive ice in greenland, it was rink’s detailed observations and descriptions that provided documentary evidence, and provided the basic background for the numerous subsequent glaciological investigations in greenland. in his attempts to understand the origin and dynamics of the inland ice, rink also ventured into many of the problems that concern the mechanics of calf-ice production. rink’s main thesis was that just as precipitation over land areas is drained by rivers, the ice streams deriving from the interior of greenland drain the snow and ice. ice streams act as ‘rivers’ in a surrounding area of ‘quiet’, dynamically less active ice (fig. 10; rink 1862). rink estimated the production of calf ice on the basis of glacier size and the quantity of ice in the fjords. on these criteria, rink (1857, 1862) recognised five ice streams (‘isstrømme’) of ‘the first order’, namely jakobshavn (69º10´n), tossukatek (69º5´n), den større kariak (70º25´n), den større kangerdlursoak (71º25´n) and upernivik (73ºn; fig. 10); the spelling used by rink is retained here. rink estimated the catchment area for each of the five ice streams to be at least c. 50 000 km2 (c. 1000 danish square miles) (rink 1875, p. 15), a clear indicabulletin 14: gsb191-indhold 04/12/07 14:36 side 18 a b c fig. 11. the margin of the inland ice at paakitsoq north of kangia; the central hill is about 80 m high. the three images (a: lithograph, h.j. rink, 1850. b: photograph, a. weidick, 1961. c: photograph, h.h. thomsen, 1987) illustrate changes in the ice margin at this site. in 1850, the gla cier front was advancing, reaching a maximum around 1880. since then a gradual thinning has taken place as can be seen from the photographs from 1961 and 1987. 19 bulletin 14: gsb191-indhold 04/12/07 14:36 side 19 20 a b c fig. 12. the lobe of the inland ice margin at paakitsoq, seen from the south-west; the central hilltop is about 150 m high. a: r.r. hammer’s drawing from 1883, which marks the maximum historical extent of the glacier lobe. the photographs from 1961 (b: a. weidick) and 1987 (c: h.h. thomsen) illustrate the progressive thinning of the ice margin. bulletin 14: gsb191-indhold 04/12/07 14:36 side 20 21 tion that he recognised the extraordinary size of the greenland ice sheet. however, while recognising the large calf-ice production from these major outlets, rink did not appreciate the unique status of jakobshavn isbræ at ilulissat. this was first recognised in the second half of the 19th century, during investigations on the rate of movement. several illustrations in rink’s papers provide details of the extent of the glacier cover in the middle of the 1800s, of which figs 11 and 12 are examples showing paakitsoq, north of ilulissat. observations and mapping of the ice margin around disko bugt as mentioned above, interest in greenland glaciers increased during the second half of the 19th century. an increasing number of scientists visited the ice margin around disko bugt, and most of these provided descriptions of the front or the inland ice margin around jakobshavn isbræ. a chronological list of the most significant visits after rink is given below. 1867. the british mountaineer edward whymper made a visit to the area around qajaa in kangia. he noted that the glacier ice was too crevassed to allow passage to the inland ice (whymper 1873; nordenskiöld 1886, p. 121). 1870. a.e. nordenskiöld visited west greenland, and described the icefjord and the front of jakobshavn isbræ. he could not determine the boundary between the front of the glacier and the calf ice in the fjord (nordenskiöld 1871; engell 1904). his visit to jakobshavn isbræ was undertaken in connection with one of the early attempts to visit the interior of greenland, and nordenskiöld also visited nordenskiöld gletscher (akuliarutsip sermerssua), c. 90 km south of jakobshavn isbræ. here, from the head of arfersiorfik fjord, he led a reconnaissance expedition that reached 56 km into the ice sheet at an altitude of 670 m a.s.l. in 1883, from the same starting point, another party led by nordenskiöld reached c. 350 km into the inland ice to an altitude of 1947 m. surface features of the ice, such as cryoconite holes (meltholes) and glacier spouts, were described. 1875. a. helland visited jakobshavn isbræ in july 1875 and conducted the first measurements of the rate of movement, which were published together with a description of the position of the front of jakobshavn isbræ (helland 1876). he also measured the movement of sermeq avannarleq in torsukattak icefjord, c. 100 km north of jakobshavn isbræ, and described the surface of the ice margin at paakitsoq. 1879. k.j.v. steenstrup (1883a) also visited sermeq avannarleq in torsukattak icefjord, and measured the rate of movement of this glacier in may 1879 and 1880. 1879. r.r.j. hammer mapped the entire fjord system around kangia, including the glaciers at the heads of the tributaries of sikuiuitsoq and tasiusaq (hammer 1883). hammer also determined the position of the front of jakobshavn isbræ in september 1879. 1880. r.r.j. hammer repeated his visit to jakobshavn isbræ in march and august 1880. a winter advance of c. 1 km and a subsequent summer recession of c. 2 km were recorded. the rate of movement was determined from the southern edge of the glacier front to its central part. hammer found that the movement was not uniform, and could find no relationship between air temperature and glacier vel ocity. the position of the front was observed to be very variable, and hammer concluded that large icebergs were released by fracturing of the ice front due to the buoyancy of the floating part of the glacier. hammer also tried to evaluate the hydrographic conditions of the icefjord. 1883. r.r.j. hammer mapped the eastern parts of disko bugt between c. 68º30´ and 70ºn (hammer 1889), and made a sketch of jakobshavn isbræ while attempting to determine the glacier recession since 1851. 1883. a.e. nordenskiöld revisited nordenskiöld gletscher (see 1870 above). 1886. robert e. peary made an attempt to reach the interior of the inland ice from a starting point at paakitsoq, 40 km north of kangia. accompanied by c. majgaard he reached a point c. 185 km into the ice sheet at an altitude of almost 2300 m a.s.l. peary also described the landscape of the ice margin (peary 1898). 1888. s. hansen made a sketch map of the front of jakobshavn isbræ for the royal danish sea chart archive (engell 1904). according to engell, a photograph from this visit showed the frontal position of jakobshavn isbræ to lie farther to the east than indicated on hansen’s 1888 map. 1891. in june 1891 the german polar explorer erich von drygalski visited the north side of kangia. 1893. von drygalski visited the area to the south of the front of jakobshavn isbræ in february 1893, and plotted the frontal position on a map, although it was difficult to determine the position precisely (von drygalski 1897). 1902. m.c. engell measured the frontal position of jakobshavn isbræ in july 1902, and showed that the recession had continued (fig. 13). velocity determinations were similar to those previously measured (helland 1876; hammer 1883; engell 1904). in addition, engell made an extensive description of the whole fjord region, and produced detailed maps of the front of jakobshavn isbræ and a b c bulletin 14: gsb191-indhold 04/12/07 14:36 side 21 22 the ice margin at the head of orpissooq fjord in the southeastern corner of disko bugt. 1903. m.c. engell again visited jakobshavn isbræ in july 1903, but could only determine the frontal position with some uncertainty; it appeared to be c. 350 m farther west than in 1902. engell observed the release of a large iceberg (engell 1910). 1904. m.c. engell made his third visit to the front of jakobshavn isbræ in the summer of 1904 although, the position of the glacier front was not determined (engell 1910). he mapped the interior of disko bugt from c. 68º 45´n (head of tasiusaq fjord) to c. 70º 05´n (torsukattak icefjord). 1912. alfred de quervain and paul-louis mercanton made a west-to-east crossing of the inland ice from eqip sermia in disko bugt to ammasalik in east greenland. a description of the ice margin around eqip sermia, c. 70 km north of ilulissat, was later published (de quervain & mercanton 1925). 1913. johan p. koch and alfred wegener made an east-towest crossing of the inland ice during their 1912–1913 expedition. in august 1913 the two scientists visited jakobshavn isbræ, and their determination of the frontal position indicated a significant recession since 1902 (koch & wegener 1930). 1929. during the 1929 preparations for the german alfred wegener expedition to the inland ice in 1930–1931, reconnaissance for an alternative route to the ice sheet was made in disko bugt. during may and june, the equipment was tested and ablation measured on a route that extended 150 km from eqip sermia north-eastwards into the inland ice reaching an altitude of 2090 m. jakobshavn isbræ was visited in september 1929, and the front position and the velocity determined. the frontal positions of eqip sermia and the glaciers in torsukattak icefjord were also described (georgi 1930; wegener et al. 1930). 1931/32. the position of the glacier front of jakobshavn isbræ was depicted on the first 1:250 000 scale map sheet of the jakobshavn area, issued by the geodetic institute, copenhagen. 1934. martin lindsay started a traverse of the inland ice from eqip sermia, from the same starting point that de quervain and mercanton used in 1912. the main objective of the three-man expedition was to explore the mountainous region south-west of scoresby sund in east greenland (fristrup 1966). 1936. eigil knuth and paul-emile victor participated in a french/swiss/danish expedition that crossed the inland ice from the ice margin c. 80 km south of ilulissat to ammasalik in east greenland. a description of the icemargin landscape was published by knuth (1937). after 1936. during and after the world war ii, there were rapid developments in aerial photography techniques. the increase of information can be illustrated by the archive of aerial photographs of the region available at the national survey and cadastre in copenhagen, dating from the years 1942, l946, 1948, 1953, 1957, 1958, 1959, 1964 and 1985. large-scale glaciological projects after world war ii after world war ii, a series of major investigations of the inland ice was carried out by military and scientific organisations. a detailed account of this work is outside the scope of this bulletin, and the investigations are only briefly mentioned here as a background for the research around disko bugt itself. reviews of the history of exploration of the greenland ice sheet and the significant results achieved are given by fristrup (1966), reeh (1989) and dansgaard (2004). the new era of scientific expeditions was initiated by the expéditions polaires francaises (epf 1948–1953), which continued the work of the german alfred wegener expedition of 1929–1931. epf worked along an east–west profile from coast to coast over the central part of greenland, and in addition to geodetic and geophysical work carried out mass-balance measurements along the route. detailed mapping of the ice sheet margin and descriptions of the area around eqip sermia were also made. the successor of epf was the expédition glaciologique internationale au groenlande (egig 1957–1960), an international collaboration between austria, denmark, france, germany and switzerland. egig included in its programme a project that aimed at photogrammetric determinations of the rate of movement and estimates of calfice production of all outlets from the inland ice that drain into disko bugt and the uummannaq fjord. vertical aerial photographs were used for this work. in 1957, flights were repeated at intervals of four to five days (bauer et al. 1968a), and in 1964 for the same glaciers at intervals of about two weeks (carbonell & bauer 1968). the project confirmed the paramount role of jakobshavn isbræ with respect to calf-ice production, compared to other productive glaciers in greenland (fig. 5). variations in the frontal position of jakobshavn isbræ were also determined (fig. 13). the egig work continued after 1960 with follow-up projects. one important task was to determine the thickness of the inland ice, and hence also the elevation and topography of the landscape below the ice, by means of airborne radar. prior to this initiative, the thickness had only bulletin 14: gsb191-indhold 04/12/07 14:36 side 22 23 been measured along a few profiles traversing the inland ice (holtzscherer & bauer 1954), using seismic, gravimetry or radar techniques. the technical university of den mark (dtu) in collaboration with the us national science foundation and the scott polar research institute in england modified the radar technique for use in aircraft. during six seasons between 1968 and 1976, more than 60 000 km of profiles were flown, and for the first time a large-scale map of the landscape under the inland ice became available (gudmandsen & jakobsen 1976). at that time, this subsurface map was a major breakthrough with significant implications for the understanding of ice dyna mics, quaternary geology and geomorphology. however, details such as the continuation of the deep fjords under the present inland ice margin could not be resolved, because radar waves could not penetrate the chaotic, crevassed ice of the ice streams. this problem was solved at jakobshavn isbræ by applying seismic methods (clarke & echelmeyer 1996). american military groups developed ways of travelling on the ice sheet, erected and maintained inland ice stations, and were also directly involved in scientific operations. the latter included the first deep drilling through the ice sheet at camp century in 1963–1966 (langway 1970; langway et al. 1985). the first systematic mapping of snow accumulation over the entire inland ice was carried out by cold regions research and engineering laboratory, corps of engineers, us army (crrel) in 1952–1955 and 1959–1960, and led to a division of the inland ice surface into dry snow, percolation facies and wet-snow facies according to altitude (benson 1959, 1962, 1994, 1996; ragle & davis 1962). more recently, the national aeronautics and space administration (nasa) has been responsible for the development and maintenance of the present satellite and airborne monitoring system (sohn et al. 1997a; williams & hall 1998; thomas et al. 2001). this system provides detailed information on current changes of surface elevation, volume and rate of movement of the greenland ice sheet. the development of the global positioning system (gps) and airborne and satellite-based altimetry has resulted in a vast expansion of data. the main focus has been on assessing the impact of climate change on the mass balance of the ice cover, especially with respect to volume changes and surface movements of the major outlets of the ice sheet (garvin & williams 1993). in 1994, nasa conceived a 0 2000 1950 1900 1850 1800 5 ye ar ( a .d .) distance (km) 10 15 20 25 30 35 40 ? 10 km 19 6419 53 19 13 19 42 19 31 19 29 m ar ch 2 00 3 50°w 69°10´n 19 02 18 79 18 83 18 9318 51 18 75 18 80 a b m ay 2 00 3 d ec 2 00 4 ju ly 2 00 5 a ug 2 00 7 a ug 2 00 6 fig. 13. a: frontal positions of jakobshavn isbræ from 1850 to the present. modified from bauer et al. (1968a) and weidick et al. (2004a). b: recessional curve of the jakobshavn isbræ glacier front, up to 1964 based on the positions given by bauer et al. (1968a). the younger parts of the curve are based on satellite information by stove et al. (1983), sohn et al. (1997a, b), later landsat, aster and modis images (alley et al. 2005; cindy starr, nasa (personal communication, 2007)) and data from a reconnaissance in 2005 (f. nielsen, personal communication 2006). the width of the curve depicts the range of seasonal variations in the positions of the glacier front. note the rapid break-up and recession from 2002. bulletin 14: gsb191-indhold 04/12/07 14:36 side 23 24 program for arctic regional assessment (parca), which for a decade has collected mass balance data covering the entire ice sheet, with a back-up of ‘ground-truth’ stations (abdalati 2001; abdalati et al. 2001). the programme includes the collection of ice-thickness data around jakobshavn isbræ (gogineni et al. 2001), where one of the ‘ground-truth’ stations is located. this is the ‘swiss station’ established at the equilibrium line near jakobshavn isbræ by the eidgenössische technische hochschule zürich, switzerland (eth), and later operated by the university of colorado (steffen & box 2001; zwally et al. 2002). with respect to subsurface mapping around jakobshavn isbræ, some details were added during investigations of the hydropower potential for the towns of ilulissat and qasigiannguit 1982–1992 (thomsen et al. 1989; braith waite 1993). these investigations collected data for the ‘quiet’ parts of the inland ice margin (the areas between the ice streams), using a radar device carried by helicopters. the device was developed by dtu and modified by the geological survey of greenland (thorning et al. 1986; thorning & hansen 1987). the most recent radar thickness measurements of the ice sheet (with particularly detailed coverage around jakobshavn isbræ) were published by bamber et al. (2001), gogineni et al. (2001) and layberry & bamber (2001). deep cores from the ice sheet the constant deposition of snow over the interior parts of the ice sheet results in the compaction of the underlying snow and conversion to glacier ice, which in the course of time flows downwards and outwards towards the margins of the inland ice. the discovery of the capability of the ice to preserve information about the climate at the time of snow deposition has given remarkable results with respect to climate change and related geological and atmospheric changes. a number of intermediate and deep ice cores have been recovered from the greenland ice sheet (reeh 1989). of the five deep cores (fig. 3), the first was made at camp century (77º11´n, 61º08´w) in 1964–1966, and had a length of 1390 m (dansgaard et al. 1969; langway 1970; langway et al. 1985; reeh 1989). subsequently, a 2037 m long core was retrieved at dye 3, where drilling was completed in 1981 (65º11´n, 43º49´w; langway et al. 1985; reeh 1989). the success of these cores was followed up by two >3000 m deep cores at summit, on the highest point of the ice sheet, namely the grip core (72º34´n, 37º37´w; johnsen et al. 1997) and the gisp2 core (72º35´n, 38º29´w; grootes et al. 1993). the fifth deep core was completed at the northgrip site (75º06´n, 42º19´w) in 2001, and reached a depth of 3001 m (dahl-jensen et al. 2002). all cores have provided detailed climatic information about the last ice age, and the cores from the central parts of the inland ice provide information on the climate up to c. 250 000 years back in time. the data obtained from the ice cores, together with detailed information about the subsurface and surface of the ice sheet, have been incorporated into models of changes of the ice sheet with time, as well as scenarios for the future development of the inland ice (fig. 14; letréguilly et al. 1991a, b; weis et al. 1996; ritz et al. 1997; huybrechts 2002; alley et al. 2005). hydropower and climatic change the energy crisis in 1973 led to a focus in greenland on the utilisation of local energy sources, as an alternative to imported oil. in greenland, hydropower was the obvious potential source of power, but an evaluation required systematic collection of hydrological and glaciological data. as part of a project carried out by the geological survey of greenland (ggu), a series of stations was erected along the inland ice margin in west greenland between c. 60º and c. 70ºn. data were collected for a hydropower project to serve the town of ilulissat, and also a project 45 km further south to serve the town of qasigiannguit. the aim 130 000 years b.p. present ice-free land 500 km a b 2000 20002500 25003000 3000 fig. 14. a: model of the surface elevation of the greenland ice sheet during the last interglacial (sangamonian/eemian) according to letréguilly et al. (1991a, b). b: modern-day surface elevation according to escher & pulvertaft (1995). bulletin 14: gsb191-indhold 04/12/07 14:36 side 24 25 of both projects was to determine the amount of melt water runoff from the ice-sheet margin. the northernmost station at paakitsoq, c. 40 km northeast of ilulissat, was started in 1982 (thomsen et al. 1988; olesen 1989). the studies covered mass-balance measurements along a line of stakes extending from the ice-sheet margin at c. 230 m a.s.l. to c. 1050 m, close to the equilibrium line (fig. 15). the studies included drilling of ice holes for measuring ice temperatures and subglacial melt. a detailed map of the subsurface of the ice margin around jakobshavn isbræ was also produced (thomsen et al. 1988; olesen 1989; weidick 1990). the data collected from paakitsoq have not yet resulted in a decision to exploit hydropower for ilulissat. the annual potential is around 72 gwh for ilulissat and c. 11 gwh for qasigiannguit (nukissiorfiit 1995). however, a hydropower plant is now in operation near nuuk in west greenland (kangerluarsunnguaq power plant, production 185 gwh/year), and two other power plants at qaqortoq/ narsaq in south greenland and at ammassalik in southeast greenland have been established. the debate on climatic change due to the increasing greenhouse effect and the increased melting of glaciers has stressed the need for data on the actual melting at the ice margin of the inland ice. the mass-balance measurements by ggu that started in 1982 in relation to hydropower, were therefore continued in 1990 in collaboration with teams from the alfred wegener institute for polar and marine research, germany (awi) and eth (thomsen et al. 1991). energy-balance measurements at the equilibrium line were made from a permanent field station on the ice, and a programme for ice-temperature measurements was also set up in a collaboration between ggu and eth. the original stake line measured from 1982 was extended from 1100 to 1600 m a.s.l. the eth programme to study climate, energy balance and the thermal regime covered the years 1990 and 1991 (ohmura et al. 1991), with participants from the institute of arctic and alpine research (instaar, university of colorado, boulder, usa). instaar runs a long-term project that aims to investigate the effects of refreezing of meltwater runoff from the inland ice, by studies of snow and ice hydrology, heat transfer, and using modelling (pfeffer et al. 1991). satellite and airborne radar and laser altimetry with large-scale coverage is now an important tool for monitoring changes in the geometry and dynamics of the ice sheet (garvin & williams 1993; thomas et al. 2001). however, ‘ground-truth’ data are still required, and the swiss station mentioned above is now part of the network of automatic weather stations that cover the inland ice (steffen & box 2001). with respect to the dynamics of jakobshavn isbræ and the adjacent margin of the ice sheet, a number of american and swiss projects have been carried out along the margin (see also the glaciology section). swiss station50°w stake sikuiuitsoq paakitsoq temperature measurement sites td 600 700 800 100 300 900 200 400 500 1000 1100 ts ts td td ts td 69°30´n 5 km fig. 15. the drainage basin at paakitsoq, north of ilulissat, where the sermeq avannarleq glacier enters sikuiuitsoq (fig. 2), showing the positions of stakes used for mass balance calculations. ts and td mark the positions of shallow and deep thermistorstrings, respectively, that were used for measuring temperatures in the ice. the swiss station marks the highest point of the glacier hydrological survey. from thomsen et al. (1991). bulletin 14: gsb191-indhold 04/12/07 14:37 side 25 26 bedrock geology precambrian crystalline rocks such as orthogneisses, granites and metavolcanic and metasedimentary rocks dominate west greenland and represent the roots of ancient fold belts. cretaceous sedimentary rocks are found on disko and western nuussuaq, where they are overlain by palaeogene basalts (fig. 16). the crystalline bedrock east of disko bugt is dominated by archaean (c. 2800 million years old (2.8 ga)) grey orthogneisses with intercalations of mica schist, amphibolite and minor ultrabasic rocks (garde & steenfelt 1999; kalsbeek & taylor 1999). in palaeoproterozoic time, a sedi mentary cover (the anap nunâ group) was deposited unconformably on this archaean basement. both basement and cover were subsequently deformed and metamorphosed during the c. 1.85 ga old nagssugtoq idian/rinkian orogenesis. the anap nunâ group is now exposed south of the torsukattak fjord, forming a curved belt of sandstone, siltstone and minor calcareous rocks (now in greenschist facies) within the reworked archaean rocks. the palaeo geology disko bugt qeqertarsuaq ilulissat disko vaigat nuussuaq 70°n 69°n 55°w 51°w53°w 71°n palaeogene basalts upper cretaceous and palaeogene sediments precambrian basement fault 50 km fig. 16. onshore (full colours) and offshore (pastel shades) bedrock geology of the disko–nuussuaq area. modified from chalmers et al. (1999) and larsen & pulvertaft (2000). altitude above 1600 m 1000–1600 m 400–1000 m 0–400 m drainage watershed shelf transverse trough bank on shelf trough-mouth fan 500 km n fig. 17. map of greenland without the present ice cover, with altitudes corrected for the present load of the inland ice (cf. fig. 4). in this reconstruction, large parts of central greenland are shown to be drained by river systems, of which one major system flows into disko bugt. the trends of the pre-glacial river systems have been reconstructed from the topography. the offshore continuation of the drainage of the interior of greenland is reflected by prominent troughs on the shelves. elevations on land according to letréguilly et al. (1991b), with offshore morphology from funder (1989) and escher & pulvertaft (1995). bulletin 14: gsb191-indhold 04/12/07 14:37 side 26 27 proterozoic orogenesis in central west greenland can be correlated with a similar event in eastern canada, and was caused by n–s collision of two archaean continents within a large canadian–greenlandic plate-tectonic system (connelly et al. 2005). several kilometres of archaean and palaeoproterozoic overburden, which formed the upper levels of the palaeoproterozoic orogen, have since been eroded away. during the mesozoic, the precambrian shield began to fragment as a consequence of plate tectonic movements. in west greenland, down-faulting and rifting parallel to the present coast took place. this led to the formation of a sedimentary basin – the predecessor of the offshore shelf area – and later to the formation of the present davis strait. sedimentation began in the middle cretaceous (pedersen & pulvertaft 1992; henriksen et al. 2000), and cretaceous and paleocene sediments are preserved between disko (qeqertarsuaq) and svartenhuk halvø (sigguup nunaa) (69–72ºn). the sediments may originally have extended both east and south of their present area of outcrop (chalmers et al. 1999; henriksen et al. 2000). they are overlain by a thick cover of basalts of paleocene and eocene age (see fig. 16) related to sea-floor spreading and the formation of the davis strait (clarke & pedersen 1976). another consequence of sea-floor spreading was the movement of greenland towards higher latitudes. thus the disko bugt region drifted from c. 59ºn to its present-day position (c. 69ºn) during the cenozoic (hurley et al. 1981). the cretaceous and palaeogene sediments found offshore and onshore show that the disko bugt region was a depositional centre for rivers draining large parts of the interior of greenland. the subsequent eruption of basalts during late paleocene and eocene times blocked and diverted the original drainage southwards, south of disko. in general, west greenland experienced considerable epeirogenic uplift during the cenozoic (henderson et al. 1981; japsen & chalmers 2000; bonow 2004; japsen et al. 2005). this uplift shaped the present-day morphology of greenland through several cycles of uplift and erosion, which in west greenland left a coastland of peneplaned precambrian bedrock bordered by marine-shelf areas with thick sequences of cenozoic sediments. characteristic features of the shelf are the large offshore, trough-mouth fans at the shelf break that developed off major fjord systems (figs 17, 18). they seem to have developed during the neogene as an extension of the drainage of the interior of greenland. the largest submarine fan of this kind is found off disko bugt, and extends so far westwards as to reach canadian offshore territory. this sediment fan appears to have a long history of formation, with the source area comprising large parts of the interior of greenland, draining westwards into the disko bugt region. the fan depth >400 m depth <400 m hellefisk moraine system palaeogene basalt escarpment hellefisk-1 100 km in ne r eg m . d yb disko banke outer egedesminde dyb 100 300 200 20 030 0 300 1200 1300 300 20 0 200 56°w 54°w 52°w 50°w58°w 52°w54°w56°w 68°n 69°n disko kangia vaigat arveprinsen ejland 69°n 70°n 70°n 68°n store hellefiskebanke fig. 18. the continental shelf offshore disko and disko bugt. bathymetry of the southern area (including egedesminde dyb) from brett & zarudzki (1979), and of the northern area (including vaigat) from chalmers et al. (1999). the hellefisk moraine system, comprising disko banke and store hellefiskebanke, is bisected by the outer egedesminde dyb. the location of the well ‘hellefisk-1’ is plotted from risum et al. (1980). inner egm dyb, inner egedesminde dyb. bulletin 14: gsb191-indhold 04/12/07 14:37 side 27 28 developed in the miocene and continued to form during the pliocene, when a base level 300 m lower than at present has been recorded (sommerhoff 1975). as noted above, rink (1862) had initiated ideas of drainage of the interior of greenland. these were subsequently elaborated by cailleux (1952), on the basis of the first systematic mapping of the subsurface of the ice sheet around 1950 (holtzscherer & bauer 1954). the subsequent more detailed mapping of the subsurface confirms the importance of the major drainage of central parts of greenland towards the disko bugt area. the drainage patterns can be traced in the transverse troughs cutting the offshore banks, the deep fjords that cut through the coastland, and the present-day ice streams that drain the inland ice. geophysical techniques are unable to localise the actual drainage channels beneath the inland ice, and the depiction of the former drainage patterns of the interior is thus provisionally constructed as a best fit of the contours of the subsurface (fig. 17). history of glaciations and interglacials cooling during the cenozoic led to increasing intensity of glaciations. the primary cause of the pliocene–pleistocene glaciations is ascribed to the periodic and quasi-periodic changes in the earth’s orbital parameters of eccentricity, obliquity and precession, leading to changes in the distribution and amount of solar energy (fig. 19; zachos et al. 2001). local climatic developments were, however, modified by variations in the topography and the trends of oceanic currents. evidence from deep-sea sediment cores indicates a marked cooling in the oligocene, with the first formation of the antarctic ice sheet more than 30 million years (30 ma) ago (fig. 19). after a relatively warm period during the middle miocene, with a temperature maximum about 15 ma ago, global cooling continued, leading to a gradual, stepwise glaciation of both the southern and later also the northern hemisphere. this is demonstrated by the occurrence of ice-rafted debris (ird) in deep-sea sediments in the northern north atlantic (thiede et al. 1998). the gradual expansion of continental ice can be traced by the oldest ird occurrence from the fram strait between north-east greenland and svalbard at 14 ma, and around iceland and south greenland about 10 ma ago (thiede et al. 1998). an ird occurrence off south greenland at 7 ma was taken as the first indication of full-scale glaciation of southern greenland (larsen et al. 1994), although this did not involve a permanent glaciation of greenland. a pliocene/pleistocene strengthening of ird pulses since c. 4 ma has been observed in the labrador sea, just south of baffin bay (thiede et al. 1998). deposition of ird in deep-sea sediments must be taken as important evidence of contemporaneous glaciations of adjacent coastal areas. however, there is still debate concerning the timing of the ird pulses and their relationship to the extent of ice cover. it has been speculated as to whether ird maxima relate to the onset/advance, or to recession/thinning/disintegration, of ice cover. discussion has also concerned the relationships between ird deposition and the type or mode of iceberg calving (reeh et al. 1999; reeh 2004). recent data from the arctic ocean (moran et al. 2006; sluis et al. 2006), however, may have implications for the glacial record summarised above. firstly, these data indicate that the eocene thermal maximum was warmer than hitherto believed. secondly, a small gneiss pebble, within sediments c. 45 ma old, is interpreted as having been icerafted, thus suggesting that the onset of glaciations was synchronous in the arctic and antarctica. onshore evidence of the first major glaciation of greenland is provided by the kap københavn formation in northern greenland. this formation includes deposits from a pre-tiglian ice age (c. 2.5 ma) that contain shell fragments from the previous warmer period (reuverian); this is succeeded by sediments referred to the subsequent warm period, the tiglian, c. 2.2–2.4 ma. the occurrence of forest tundra at this locality at this time indicates a greenland without an extensive central ice sheet (bennike 1990; funder et al. 2001). the subsequent glaciations (ice ages) grew in intensity and it is possible that the present inland ice first formed during the middle and late pleistocene (as late as c. 0.8 ma). this required sufficient cooling during the ice ages to form an ice cover that was large enough to survive melting during the intervening interglacials. in the vicinity of disko bugt, the oldest known quaternary sediments are the deposits at pattorfik (fig. 20). they were observed by k.l. giesecke and h.j. rink in the first half of the 1800s, and described and investigated in detail by símonarson (1981) and funder & símonarson (1984). amino acid analyses of shells from these marine deposits indicate that they belong to early pleistocene (1 ma or more). although situated on the southern shore of uummannaq fjord, the deposits have been protected from subsequent glacial erosion during the ice ages by a layer of lithified talus breccia. other interglacial and interstadial deposits have been discovered subsequently. these pre-holocene occurrences are nearly all located at the extreme western coastal parts of central west greenland (fig. 20). they occur from western disko, over the tip of the nuussuaq peninsula to svartenhuk bulletin 14: gsb191-indhold 04/12/07 14:37 side 28 29 halvø, and are referred to a border zone of the ice age glaciations. at present, about 25 occurrences are known between disko bugt and southern svartenhuk halvø. on the basis of faunal compositions and amino acid analyses, the deposits have been referred to four pre-holocene marine events: the interglacial ivnaarssuit marine event (early– middle pleistocene), the interglacial nordre laksebugt marine event (mid-middle pleistocene), the interstadial laksebugt marine event (middle–late pleistocene), and the last interglacial svartenhuk marine event (eemian/ sangamonian; 130 000 – 115 000 years before present (130–115 ka b.p.); bennike et al. 1994). one of the interglacial sites is located on eastern disko island (fig. 20), and consists of reworked shells in moraine (funder & símonarson 1984). during the intervening ice ages, the ice sheet expanded out to the offshore banks, implying that there was little coastal lowland on which deposits older than the holocene could be preserved. based on weathering differences, it has been established that an extensive glaciation (hellefisk glacial event) resulted in the expansion of the ice out to the shelf break south of c. 68ºn (kelly 1985); this event has been tentatively dated to the illinoian/saalian (c. 380 to 130 ka b.p., gibbard et al. 2005, or c. 300 to 130 ka b.p., geyh & müller 2005). the moraine system referred to the helle fisk glacial event is depicted in fig. 18 after brett & zarudzki (1979) and kelly (1985). the moraines are situ ated near the shelf break on the southern store helle fiskebanke, and on the central parts of disko banke, presumably because of the greater depth of this bank (fig. 18). 80 partial or ephemeral ice sheet paleocene eruptions: original drainage from interior of greenland to disko bugt area barred by plateau basalts 4 eocene thermal maximum initiation of major westward-flowing fluvial system draining central part of greenland late eocene and oligocene offshore sediment action reflects rejuvenation of relief by uplift small ephemeral antarctic ice sheets permanent antarctic ice sheet, first ice-rafted detritus in north atlantic mid-miocene thermal maximum a nt ar ct ic ic e sh ee tsn or th er n h em is ph er e ic e sh ee ts present shape and elevation of greenland notes full-scale permanent ice sheet pl io eo ce ne o lig oc en e m io ce ne pa le oc en e δ o (‰)18 qua 12 temperature (°c) pa la eo ge ne m es o c en oz oi c n eo ge ne c re t ? 5 4 3 2 01 a ge ( m a) 10 20 0 30 40 50 60 70 fig. 19. global cooling during the cenozoic is illustrated by a temperature curve based on deep-sea oxygen isotope records (from zachos et al. 2001); the temperature evolution is compared with global ice-sheet development and selected local events in greenland. cret, cretaceous; meso, mesozoic; plio, pliocene; qua, quaternary. bulletin 14: gsb191-indhold 04/12/07 14:37 side 29 30 there is little information to constrain the extent of the ice-age ice sheets of west greenland. as mentioned above, it is possible that the illinoian/saalian ice cover extended over store hellefiskebanke out to the shelf break in the shelf areas south of disko island. the till on store hellefiskebanke covers a thick suite of eocene and younger deposits (whittaker 1996; chalmers et al. 1999). in contrast, on the eastern parts of disko banke, a ‘rough sea floor type’ seems to indicate that volcanic rocks underlie the quaternary sediments (brett & zarudzki 1979). early investigations reported erratic gneiss boulders on the basalt terrain of disko, and on the outer parts of nuussuaq and svartenhuk halvø farther to the north (steenstrup 1883b). on qilertinnguit on the north coast of nuussuaq (fig. 1), steenstrup reported erratic occurrences up to 1200 m a.s.l., and in nordfjord (kangersooq) and around disko fjord (kangerluk) on disko he noted occurrences up to 920 m and 628 m a.s.l. respectively (for location see fig. 2). all three localities are situated close to the outer coast and could be related to the illinoian glaciation, or even older ice ages. if so, the outer egedesminde dyb (fig. 18) must have repeatedly served as a southern drainage outlet from disko bugt for a high-arctic ice stream, probably moderate in size and with a depth of only 600 m. the calf-ice production was presumably also moderate, as can be seen from the relationship between the mode of calving of ice sheet outlets and temperature conditions (reeh 1994; reeh et al. 1999). the last interglacial (sangamonian/eemian) and the last ice age (wisconsinan/weichselian) it was noted above that the hellefisk glacial event may have been followed by the svartenhuk marine event, which is referred to the last interglacial stage. the sang amo nian/ eemian interglacial has been described as more humid and warmer than the present interglacial, with a mean summer temperature up to 5ºc higher than in the holocene (bennike & böcher 1994). based on the climate record from ice cores and the altitude of the surface and subsurface of the present ice sheet, a detailed picture of the changes in the extent of the ice sheet throughout the last 130 000 years has been modelled (letréguilly et al. 1991a, b; weis et al. 1996). these models indicate that ice-sheet recession during the eemian interglacial was so extensive that the inland ice almost split into a large northern ice sheet and a minor southern one, and left large parts of south-western greenland ice-free (fig. 14). the stronger recession in the south-west is ascribed to the generally low elevation of this region, which together with strong ablation resulted in a rapid recession. loss by calving through ice streams may at this time have been confined to outlets north of disko bugt. an extensive reduction of the inland ice is also indicated by the character of the basal ice in the camp century and dye 3 ice cores (koerner 1989). the basal silty ice from the dye 3 core has tentatively been dated at 450–800 ka b.p., which suggests that the dye 3 region was not de glaciated during the last interglacial (willerslev et al. 2007). disko bugt uummannaq fjord n inland ice godhavn stade/ice shelf? ea rl y pl ei st oc en e m id dl e pl ei st oc en e la te pl ei st oc en e h ol oc en e pátorfik me svartenhuk me laksebugt me nordre laksebugt me ivnaarssuit me/mudderbugt az tiglian illinoian/saalian wisconsinan/weichselian sangamonian/eemian q ua te rn ar y age local stratigraphy 50 km 780 130 11.7 a ge ( ka b .p .) 2600 nuussuaq fig. 20. stratigraphic summary of interglacial and interstadial deposits between disko bugt and southern svartenhuk halvø, based on kelly (1986) and bennike et al. (1994). az, aminozone; me, marine event. the accompanying map shows the location of these deposits; the symbols represent one or more localities. bulletin 14: gsb191-indhold 04/12/07 14:37 side 30 31 the subsequent build-up of the inland ice has been modelled by letréguilly et al. (1991b). the general trend shows a slow temperature decrease from the end of the eemian through the early and middle parts of the wisconsinan, followed by large oscillations as shown by the ice-core records (dansgaard et al. 1993). during shortlived cold periods at about 70 ka b.p. and 25–22 ka b.p., the temperature fell to around 23–25ºc below the present (johnsen et al. 1995; dahl-jensen et al. 1998; dansgaard 2004). throughout the wisconsinan, the ice-core records show abrupt temperature changes of 10–12ºc that seem to be related to abrupt changes in the course and intensity of ocean currents (broecker et al. 1985). the last glacial temperature minimum has been dated to around 21.5 ka b.p. (johnsen et al. 1995). around disko bugt, as elsewhere in greenland, the last glacial maximum (lgm) can be related to offshore features and deposits, although the exact position of the ice margin during the lgm is still a matter of debate. the ice margin at the lgm is usually related to offshore morphological features of the banks and the intervening troughs (‘dyb’, the transverse channels of holtedahl 1970). in central west greenland, the most prominent of these troughs is the c. 350 km long egedesminde dyb – where an iceage predecessor of jakobshavn isbræ may have been located (figs 18, 21). the south side of egedesminde dyb is limited by store hellefiskebanke, one of the largest of the offshore banks in west greenland. at the most shallow locality, the water depth is only 8 m. an exploration well (hellefisk-1; fig. 18) drilled on the western slope of the bank revealed a 200 m thick cover of probable glacial till, overlying c. 3 km of cenozoic sediments resting on basaltic lava flows (risum et al. 1980). geophysical profiling has also revealed thick deposits of cenozoic sediments (henriksen et al. 2000). the surface morphology indicates significant marginal moraines, of which a western system near the shelf break (as noted above) could be referred to the illinoian. another system of moraines is found around the eastern slopes of store hellefiskebanke, and on the banks farther south. these 100 km none moderate intense iceberg scouring store hellefiskebanke in ne r eg m . d yb disko banke outer egedesminde dyb 100 300 200 20 030 0 300 1200 1300 300 20 0 200 70°n 69°n 68°n 50°w58°w 56°w 52°w 67°n 54°w fig. 21. distribution and intensity of iceberg scouring on the shelf offshore west greenland, from brett & zarudzki (1979). inner egm dyb, inner egedesminde dyb. bathymetric contour intervals are 100 m. bulletin 14: gsb191-indhold 04/12/07 14:37 side 31 32 moraines continue as marginal and terminal moraine systems surrounding the troughs between the banks, and their lobate nature could indicate that the moraines were originally deposited on land (funder 1989). during the wisconsinan ice age, the global sea level was about 130 m below the present (lambeck & chappell 2001). however, the increased glacier load in greenland during this period would have reduced this figure, and abrasion terraces and beach ridges down to 70 m below present sea level in west greenland imply that the wisconsinan ice margin rested on dry land to a large degree (sommerhoff 1975). the ‘inner moraines’ of the offshore areas have been correlated with the oldest and highest situated moraines of the coastal high mountains (distinguished by their degree of weathering) and with the period described as the ‘sisimiut glaciation’ (kelly 1985). disko banke, located south-west of disko, is bounded to the south-east and south by the troughs referred to as the inner and outer egedesminde dyb (fig. 18). the shallowest part of disko banke has a water depth of c. 150 m, with the surface only partly till-covered. eocene and younger sediments dominate the western part of disko banke (chalmers et al. 1999), whereas the surface of the eastern part comprises volcanic rock ridges with a thin, discontinuous layer of till (zarudzki 1980). the basalts dip gently westward, with steep scarps to the east (brett & zarudzki 1979). a pronounced east-facing escarpment connects disko banke in the north with the easternmost part of store hellefiskebanke in the south. this escarpment forms a threshold with depths of 200–300 m below present sea level. it divides egedesminde dyb into an eastern, deep and narrow channel with maximum depths over 1000 m (inner egedesminde dyb) and a shallower western part with depths up to around 600 m (outer egedesminde dyb; fig. 18). if the moraine system on store hellefiskebanke does mark the outer limit of an extensive illinoian glaciation, then the ‘reduced’ extent of a wisconsinan glaciation implies that the western shores of disko were only characterised by a shelf glaciation (bennike et al. 1994). along the coast of southern disko, near the town of godhavn, are major moraines that were formed during the godhavn stade (fig. 22; ingólfsson et al. 1990). these moraines are taken as evidence for the maximum extent (lgm) of the greenland ice sheet during the wisconsinan. the main moraine, ‘pjetursson’s moraine’, is c. 1.5 km long and reaches an elevation up to 220 m a.s.l. east of godhavn. smaller moraines are found up to 110 m a.s.l. about 15 km farther to the west. although closely related to the marine limit of the area, it cannot be excluded that the godhavn stade marks an early re-advance, or halt, during recession from the lgm. 70°n 69°n 54°w t t g g d m t m 12.4 11.8 11.6 10.3 9.9 10.0 10.5 es ca rp m en t 10.0 9.8 9.8 10.5 9.6 10.0 9.3 9.9 6.18.4 25 km 7.3 7.7 8.4 7.9 jakobshavn isbræ kangilerngata sermia tors ukattak eqip sermia disko bugt disko kangia 52°w vaigat nuussuaq akulliit 50°w nordenskiöld gletscher fig. 22. radiocarbon dates pertaining to the last deglaciation of the ice-free areas around disko bugt. dates are given in calibrated thousand years before present (cal. ka b.p.), using the intcal04 data set for calibration (reimer et al. 2004). based on dates published by ingólfsson et al. (1990), rasch (1997), bennike et al. (1994), bennike & björck (2002), long & roberts (2003), long et al. (2006) and this study (appendix 2); further details are given in tables 2 and 3. the map also shows the approximate trend of the basalt escarpment west of the mouth of disko bugt, the marrait moraine system (m), the tasiussaq moraine system (t), the drygalski moraines (d) and moraines of the godhavn stade (g). bulletin 14: gsb191-indhold 04/12/07 14:37 side 32 33 the outer egedesminde dyb was described as a valley by zarudzki (1980), who also noted that “well-preserved flank moraines, found at gradually lower elevations in the valley, testify to a recent withdrawal of the ice” (zarudzki 1980, p. 60). it may be suggested that the outer egedesminde dyb drained a high-arctic type ice stream during the lgm, as well as during earlier more extensive glaciations, although it was probably restricted in size and also in calf-ice production. the northern drainage route for wisconsinan ice, the vaigat (sullorsuaq) strait, is a typical glaciated fjord with depths over 600 m in its south-eastern part and around 200–300 m near hareøen (qeqertarsuatsiaq) at its western mouth. quaternary deposits with a thickness of several hundred metres are found in the strait (denham 1974). there is general agreement on a ‘reduced’ extent of the wisconsinan ice sheet during the lgm. the margin of the ice sheet may have been situated on the proximal parts of the offshore banks south of disko bugt, with outlets in the intervening transverse troughs reaching terminal moraines (or thresholds) near the shelf break. the northern conduit of ice from disko bugt through vaigat only filled the strait as far as a position near hareøen at its mouth. this is in accordance with the concept of a lgm limit at the mouth of the fjords farther north in west greenland (funder 1989). the acceptance of the godhavn stade at or near the lgm on southern disko near ege desminde dyb, would leave the whole of western disko unglaciated by the ice sheet, but presumably with local glaciers of high-arctic type contributing to shelf ice in baffin bay. the nature of this shelf may have been comparable to that of the present ellesmere island ice shelf in canada, which has a thickness of up to 100 m, and shows a gradual transition between an ice shelf evolved from sea ice and a shelf of glacier ice (jackson 1997; jeffries 2002), or to the ice shelves in northern greenland (koch 1928; higgins 1989). ice shelves are known to be particularly sensitive to climatic changes. this is well documented for ice shelves in antarctica, such as the ross shelf (bindschadler & bentley 2002), and also for the ice shelves of northern ellesmere island adjacent to the arctic ocean (jeffries 2002). a comparable great variability in ice-shelf extent, related to former climatic oscillations, would be expected during ice ages. a record of extensive iceberg scouring of the seabed in the eastern part of disko banke and the northern slope of store hellefiskebanke (fig. 21; brett & zarudzki 1979), points to the close proximity of a calving glacier margin at disko banke; the higher, northern part of store helle fiskebanke may have been dry land. iceberg scouring is recorded to a depth of 340 m, and it is assumed that the 300 m deep bedrock threshold across the outer and inner part of egedesminde dyb did not allow icebergs with a greater draft to pass. it is presumed that the deeper scours were created by icebergs calved in outer egedesminde dyb when the ice extended beyond the threshold. the relative sea level at the end of the lgm may also have been higher than at present. the escarpment at egedesminde dyb south of disko may have had a braking effect on the glacier ice, but the width of the glacier (c. 130 km) may have compensated for the reduced ice thickness. the subsequent concentration of ice locality latitude longitude laboratory material 14c age† (± 2σ) rcorr14c age§ calib. age¢ mean age reference n w number* years b.p. years b.p. years b.p. ka b.p. hareøen 70°23´ 54°57´ ua-1789 shells 10870 ± 130 10470 12800 – 11950 12.4 bennike et al. (1994) w nuussuaq 70°28´ 54°02´ aar-3496 bryophytes 10160 ± 75 12100 – 11400 11.8 bennike (2000) nw disko 70°16´ 54°37´ i-16393 shells 9920 ± 150 9920 12050 – 11050 11.6 bennike et al. (1994) e nuussuaq 70°04´ 52°06´ k-994 shells 8940 ± 170 8940 10500 – 9550 10.0 weidick (1968) e disko 69°40´ 52°01´ k-3667 gyttja 8950 ± 125 10400 – 9600 10.0 ingólfsson et al. (1990) e disko 69°40´ 52°00´ k-3660 shells 8700 ± 120 8700 10200 – 9500 9.9 ingólfsson et al. (1990) arveprinsen ejland 69°46´ 51°15´ beta-107879 gyttja 8820 ± 100 10200 – 9550 9.9 long et al. (1999) godhavn 69°17´ 53°28´ aar-5 shells 9650 ± 250 9250 11250 – 9750 10.5 ingólfsson et al. (1990) central disko bugt 69°11´ 51°49´ aa-37711 foraminifers 9483 ± 65 9083 10500 – 10150 10.3 lloyd et al. (2005) egedesminde 68°36´ 52°34´ hel-362 shells 8970 ± 170 8970 10500 – 9550 10.0 donner & jungner (1975) s egedesminde 68°26´ 52°57´ aa-38842 gyttja 9330 ± 99 10800 – 10200 10.5 long & roberts (2003) se disko bugt 68°40´ 51°07´ aa-39659 gyttja 8585 ± 86 9790 – 9430 9.6 long & roberts (2002) table 2. radiocarbon age determinations relating to the last deglaciation (see also table 3) * sources of age data: ua: ångström laboratory, uppsala. aar: aarhus ams c14 dating centre. i: teledyne isotopes. k: the former radiocarbon laboratory in copenhagen. beta: beta analytic. aa: the nsf-arizona facility. hel: the dating laboratory at university of helsinki. † the radiocarbon age determinations from ua and aar have been corrected for isotopic composition by normalising to –25‰ on the pdb scale, and those from k by normalising to 0‰. the date from i has not been normalised. § rcorr: reservoir-corrected. the age determinations on marine material from ua and aar have been seawater reservoir corrected by subtracting 400 years. the dates from k and i have not been corrected (bennike 1997). ¢ calibrated using the intcal04 data set (reimer et al. 2004) and the oxcal. v.3.10 software program (bronk ramsey 2001). bulletin 14: gsb191-indhold 04/12/07 14:37 side 33 34 masses through the funnel-like shallow conduits in the bedrock of the eastern part of disko banke might then have led to a concentration of ice in the outer egedesminde dyb. the high-arctic conditions of the environment and the ice cover at that time (low precipitation, low temperature) suggest that this outer egedesminde dyb ice stream was less important than the present jakobshavn isbræ; it may have been similar to present glaciers in northern greenland (reeh 2004) or antarctica (thomas 1979; swithinbank 1988; jacobs et al. 1992). these glaciers, or ice streams, are characterised by greater width and lower velocities than present ice streams in the southern part of greenland (rignot & kanagaratnam 2006). in addition, they have long, floating tongues with a high rate of bottom melting. calf-ice production, precipitation and ablation are more limited than in present-day ice streams in the southern parts of greenland. the dynamic differences between arctic and antarctic ice streams are considered by truffer & echelmeyer (2003). the oldest radiocarbon age determination from the mouth of disko bugt is 10.5 ka b.p. (table 2; ingólfsson et al. 1990), which post-dates the godhavn stade mentioned above (fig. 22). for the north-western entrance to the vaigat strait, the information is sparse, but a shell has been dated to 12.4 ka b.p. (bennike et al. 1994; bennike & björck 2002); an estimate of the extent of the ice cover before that time can only be speculative. to summarise the events from the lgm (22 ka b.p.) to the glacial situation at 13–10 ka b.p., it is suggested that onset of a climatic initial warmth around 20 ka b.p. was followed at c. 19 ka b.p. by a rise in sea level, with a breakup of the outer parts of marine ice shelves and margins at around 18 ka b.p. at 16–14.5 ka b.p., a further rise in sea level coupled with the allerød/bølling warm period (14.7–12.6 ka b.p.; lambeck & chappell 2001) accelerated the process of thinning of the ice margin and breakup of the ice shelf. during the younger dryas (12.6–11.7 ka b.p.) the ice margin may have receded to a position near the basalt escarpment between disko banke in the north and store hellefiskebanke in the south. this escarpment must have formed a barrier that retained and hampered the glacier ice flow to the sea. thus little recession of the ice margin took place. the vaigat lobe of the ice sheet north of disko probably reached the outer part of the vaigat strait. from the extent of the ice margin during the godhavn stade, and from comparisons with the present-day surface profile of the ice sheet, it is estimated that the ice cover over disko bugt had a thickness of 1000–1500 m. this estimate is in agreement with the elevation of the nunatak moraines on the outer coastal highland farther south around 67ºn, the ‘taserqat stade’ of weidick (1972a). locality latitude longitude altitude laboratory material£ 14c age† (± 2σ) rcorr14c age§ calib. age¢ δ13c reference n w m a.s.l. number* years b.p. years b.p. ka b.p. ‰ ilulissat and qasigiannguit areas sermermiut 69°12´ 51°04´ 1–2 ua-1086 mc shell 8795 ± 130 8395 9.3 0 ¶ this study ‘qassortoq’ 69°06´ 51°04´ 5–10 k-1818 shells 8630 ± 130 8630 9.8 weidick (1972b) pinguarssuit 69°05´ 51°08´ 21 i-6243 shells 6835 ± 125 6835 7.7 weidick (1973) ‘sandbugten’ 69°03´ 51°08´ ? k-2022 shells 7690 ± 120 7690 8.6 weidick (1974a) lersletten 69°02´ 51°01´ 36 k-992 shells 7110 ± 140 7110 7.9 weidick (1968) narsarsuaq 69°02´ 51°01´ 33 k-987 gyttja 7850 ± 190 8.8 tauber (1968) tasiusaq 69°02´ 50°56´ 10–15 ua-4575 pa shell 8140 ± 95 7740 8.7 –14.49 this study ‘lerbugten‘ 69°01´ 51°08´ ? k-2023 shells 8680 ± 135 8680 9.8 weidick (1974a) marraq, w. part 69°00´ 51°07´ 25 ua-4574 shell 9180 ± 75 8780 9.9 –0.32 this study eqaluit 68°56´ 50°58´ 25 k-993 shells 7650 ± 140 7650 8.5 weidick (1968) eqaluit 68°56´ 50°58´ 20? ua-4573 mt shell 8215 ± 80 7815 8.7 1.21 this study serfarsuit 68°50.5´ 50°47´ 15 ua-4572 mt shell 7500 ± 75 7100 7.9 1.33 this study other areas ussuit 67°51´ 50°16´ 42 k-1556 shells 6760 ± 130 6760 7.6 kelly (1973) eqip sermia 69°46´ 50°13´ 0.8–2.3 k-6373 shells 6420 ± 110 6420 7.3 1.4 rasch (1997) qapiarfiit 69°52´ 50°19´ 2–3 k-3663 shells 7600 ± 110 7600 8.4 –0.1 ingólfsson et al. (1990) table 3. selected early holocene radiocarbon age determinations (see also table 2, appendix 2) * sources of age data: ua: ångström laboratory, uppsala. i: teledyne isotopes. k: the former radiocarbon laboratory in copenhagen. £ mc: macoma calcarea, pa: portlandia arctica, mt: mya truncata. † the radiocarbon age determinations from ua have been corrected for isotopic composition by normalising to –25‰ on the pdb scale, and those from k by normalising to 0‰. the date from i has not been normalised. § rcorr: reservoir-corrected. the age determinations on marine material from ua have been seawater reservoir corrected by subtracting 400 years. the age data from k and i have not been corrected (bennike 1997). ¢ calibrated using the intcal04 dataset (reimer et al. 2004) and the oxcal. v.3.10 software program (bronk ramsey 2001). ¶ assumed value. bulletin 14: gsb191-indhold 04/12/07 14:37 side 34 35 the collapse of the ice cover in disko bugt the seabed of the southern and central parts of disko bugt is mainly 200–400 m below sea level, and characterised by a rugged bedrock terrain (fig. 18; brett & zarudzki 1979). seismic data reveal quaternary deposits of 100 m or more in some places (denham 1974; chalmers et al. 1999). on the bathymetric map (fig. 18), a submarine divide connects eastern disko and arveprinsen ejland (alluttoq). from this area, depressions (‘drowned glacial valleys’) lead either north-westwards to the vaigat strait or south-westwards towards egedesminde dyb. off kangia, two e–w-trending channels can be seen as a continuation of the fjord (long & roberts 2003), although they are only about 400–500 m deep. minimum ages for the chronology of the deglaciation of disko bugt are provided by dates of shells from raised marine deposits, from marine sediment cores or from dating of basal gyttja in lakes (fig. 22). sediment cores have been retrieved from the inner egedesminde dyb, 20–30 km east of the threshold, but only midto late holocene sediments were penetrated (kuijpers et al. 2001; jensen 2003). other cores have been collected in eastern disko bugt (lloyd et al. 2005), and a minimum date for the deglaciation here is 10.3 ka b.p. the onset of a branch of the west greenland current into disko bugt has been dated at c. 9.2 ka b.p. – at a time when jakobshavn isbræ terminated at isfjeldsbanken; it receded from this position at 7.9 ka b.p., an event indicated by a reduced sedi mentation rate seen in the cores (lloyd et al. 2005). recent detailed investigations of isolation basins, relative sea-level changes and deglaciation history have been carried out by long et al. (1999, 2003) and long & roberts (2002, 2003). as mentioned above, a minimum date for the deglaciation of the godhavn area is recorded by a date of c. 10.5 ka b.p. a date of c. 9.6 ka b.p. for basal gyttja from a lake above the marine limit provides a minimum age for deglaciation of the south-eastern corner of disko bugt (table 2; long & roberts 2002). from the area between qasigiannguit and ilulissat, several dates have been obtained on shells from basal marine deposits, described by laursen (1944, 1950) and laursen (in: weidick 1974a); selected dates are presented in table 3. the ages, between 9.9 and 9.3 ka b.p. are presumed to be related to a marine limit of c. 70 m a.s.l. from arveprinsen ejland, farther north in disko bugt, basal gyttja from isolation basins has yielded ages up to c. 9.9 ka b.p. (long et al. 1999), also related to a marine limit of about the same altitude. a dating of shells from easternmost disko also yielded an age of 10.0 ka b.p. ( table 2; ingólfsson et al. 1990). at about the same time, the vaigat strait became progressively ice-free. an age determination of 12.4 ka b.p. provides a minimum age for the deglaciation of the mouth of vaigat (bennike et al. 1994). the outer part of vaigat was deglaciated before 11.8 ka b.p. (bennike 2000) and the inner part before 10.0 ka b.p. (table 2; weidick 1968). for both routes from disko bugt to davis strait, minimum dates for the last recession are thus available, but a few details about the processes during the recession are known. the broad nature of the mouth of the bay south of disko may have led to fast recession, reinforced by a temporary ice stream in the inner egedesminde dyb (long & roberts 2003). however, the threshold between disko and store hellefiskebanke may have acted as an iceberg bank, in the same way as the present isfjeldsbanken at ilulissat restrains the icebergs coming from the present jakobshavn isbræ. halts in the recession at ‘pinning points’ at the mouth as well as in the bay (long & roberts 2003) may have influenced the rate of recession, but these halts may have been relatively short stops of decades during a fast recession. the shallow-water belt between godhavn and aasiaat, including the islands in the mouth of disko bugt, may have caused a short halt in recession. calculations of the calf-ice production of the former ice streams in disko bugt, based on an empirical correlation between calf-ice production and contemporaneous suggested water depth at calving glacier fronts (pelto & warren 1991), and combined with the ‘jakobshavn effect’, are provided by long & roberts (2003). this is an interesting approach to understanding the life of ice streams. the ‘jakobshavn effect’ is caused by rising temperatures, which lead to increased surface melting at the ice margin (fastook & hughes 1994). the meltwater drains into crevasses and moulins, warming the ice and lubricating the bed, and leading to higher velocities and increased crevassing of the surface, which again leads to increased heat transport from the surface to the bottom of the glacier. however, this is just one element in the complex interplay of mass-balance changes and glacier response to climatic change through the dynamics of the ice margin. further elements need to be included to explain the onset and demise of the individual ice streams, which are actors in the break-up of such a large segment of the ice-sheet margin as the former marine outlet covering disko bugt. another climatic element may explain the ‘disko stade’ of disko, during which local glaciers filled most of the broad valleys on eastern disko. it has been dated to about 10.7 ka b.p. (ingólfsson et al. 1990). the local readvance at this relatively late and warm time has been explained by changes in the prevailing wind systems causing heavier bulletin 14: gsb191-indhold 04/12/07 14:37 side 35 36 snowdrift from the west to the east. moraines referred to the disko stade are also found on other parts of disko, with correlation made by determination of the depression of the glaciation level. however, the violent expansion of the local glaciers during the godhavn stade in eastern disko may perhaps be related to surging behaviour. at the present day, surging glaciers are common and widespread on disko and in central east greenland, where the bedrock is dominated by basalt (weidick 1988). in recent studies, it was found that 75 out of 247 local glaciers on disko could be classified as surge-type glaciers, and that the quiescent phase could be as long as 100 years or more (yde & knudsen 2005, in press). in the broad valleys on eastern disko, large areas of relict glacier ice are seen, which can be explained by glaciers that have surged in the past. the relict ice on disko has not been dated, but neoglacial relict allerød –20 0 bølling high-arctic ice stream in the outer egedesminde dyb –10 younger dryas shelf ice and marine ice margin rising sea level causes beginning break-up of ice shelves onset of initial warmth allerød/bølling warmth accelerates process of ice-shelf break-up main iceberg scouring around the outer egedesminde dyb? halt of recession of ice margin at main sill godhavn stade? initiation of recession of the vaigat lobe activation of the inner egedesminde dyb ice stream thinning of the disko bugt ice-sheet lobe west greenland current (wgc) penetrates as far north as today break-up of the ice cover over disko bugt, formation of marrait moraine system onset of branch of wgc into disko bugt cold 8.2 ka b.p. event recession of ice margin in central disko bugt. attainment of present position at 6–7 ka b.p. at jakobshavn isbræ and at c. 8 ka b.p. in northern and southern disko bugt locally marked readvance of ice margin (narssarssuaq stade in south greenland) (drygalski stade of nuussuaq?) medieval warmth little ice age moraines last glacial maximum a ge ( ka b .p .) 0 2 4 6 8 10 12 14 16 18 20 22 temperature (°c) fig. 23. trend of the temperature development since the last glacial maximum, compiled from ice-core records (dansgaard et al. 1984; dahl-jensen et al. 1998; dansgaard 2004) and events related to the recession of the ice-sheet margin around disko bugt. bulletin 14: gsb191-indhold 04/12/07 14:37 side 36 37 ice in sanddalen, jökelbugten, north-east greenland, shows similarities to the relict ice on disko (bennike & weidick 2001). the abrupt temperature rise at the transition from the younger dryas to the holocene at 11.7 ka b.p. (fig. 23) must have led to intense melting and thinning of the ice margin, especially in a lowland or marine environment such as disko bugt. the opening of the ice-filled nearby davis strait would have led to increasing humidity and ablation. heat transfer from the ocean to the over 100 km long glacier front in western disko bugt must have been appreciable, as the west greenland current at 10.2 ka b.p. seems to have penetrated just as far north then as it does at the present day (funder 1990, 1994). an ‘ice stream’ occupying the inner egedesminde dyb may have favoured the whole process of fast disintegration of the glacier lobe in disko bugt. with its ne–sw orientation and its form, the inner egedesminde dyb cannot be classified as a typical uvalley, but its large depth could be due to glacial erosion. perhaps the inner egedesminde dyb is related to the n–soriented marginal channels, which separate the gneiss terrain near the coast from the offshore banks (holtedahl 1970). the marginal channels were formed by faulting followed by fluvial and glacial erosion. the role of the inner egedesminde dyb was gradually reduced as the ice margin receded to the north and east beyond the inner egedesminde dyb. the fast recession came to a halt when the front of jakobshavn isbræ became anchored at isfjeldsbanken near ilulissat at 9.9 ka b.p. the frontal area was, however, strongly reduced, scarcely allowing for more than a fraction of present day calf-ice production (weidick 1994a, b; long & roberts 2003). the marked increase in temperature at the transition from the younger dryas to the early holocene, at c. 11.7–10.0 ka b.p., was followed by a marked rise of global sea level. a relative sea level of c. 70 m above the present was reached at c. 10 ka b.p. in eastern disko bugt. in the disko bugt region, the recession of the glacier lobe from the escarpment began in this period, so that the outer parts of disko bugt quickly became free of glacier ice. the main thinning of the disko bugt piedmont lobe took place in the millennium after the end of the younger dryas, and at around 10.5 ka b.p. the godhavn area and areas near aasiaat were ice-free (donner & jungner 1975; ingólfsson et al. 1990; long et al. 2003). deglaciation dates show that the outer parts of disko bugt were already icefree at around 10.3 ka b.p. (fig. 22). the final break-up of the ice cover in the bay may have taken place over a few centuries or less. during the recession, at around 10.0–9.5 ka b.p., a major change in the condition of the ice margin occurred. large parts of the ice margin were now resting on land. the recession (or break-up of the marine ice) probably con tinued through the torsukattak fjord system in the northeastern part of disko bugt, or was brought to a halt in the eastern parts of disko bugt. palaeoceanographic investigations show that a strengthening of the west greenland current took place at 9.2 ka b.p. in disko bugt (lloyd et al. 2005). at this time, the front of jakobshavn isbræ had receded somewhat into kangia, as shown by the occurrence of shells at sermermiut that yield an age of 9.3 ka b.p. (table 3). the ‘fjord stage’ and the attainment of the present ice-margin position after the recession from the marine-based ice margin in disko bugt to the uplands of the islands and peninsulas in the eastern interior part of disko bugt, the morphological environment of the ice margin changed drastically. calving became restricted to a few fjords, and large parts of the ice margin became land based where ice loss was mainly by superficial melting in the ablation zone (weidick 1985). a group of moraines, kame terraces and other ice contact features were identified during reconnaissance mapping of west greenland quaternary deposits in the late 1960s and 1970s (figs 22, 24). these generally n–s-trending ice margin deposits occur locally throughout the inner part of the greenland coastland from 64º to 70ºn. it was already clear from the first description (weidick 1968), that the deposits of this former ice-margin zone were formed over a longer period, and distinction was made between an older ‘marrait moraine system’ and a younger ‘tasiussaq moraine system’ (kelly 1985). the older system formed contemporaneously with a local marine limit of c. 75 m, whereas the younger system formed when the relative sea level was c. 40 m a.s.l. a group of dates in the south-eastern corner of disko bugt (fig. 22) have been obtained from basal gyttja deposits, mainly from isolation basins, but including one lake situated above the marine limit (long & roberts 2002). it appears that the island of akulliit was ice-free before c. 9.6 ka b.p., whereas a moraine on the nuuk peninsula, 5 km further to the east-north-east, was formed around 8 ka b.p., perhaps related to the 8.2 ka b.p. cold event (fig. 23). the radiocarbon age of 9.6 ka b.p. from se disko bugt provides a minimum date for the local deglaciation, but this age is not related to ice-margin deposits (table 2). however, c. 30 km farther to the north, four ages of 9.9–9.3 ka b.p. have been obtained on shells from cliffs in basal marine silt along the shores between the narsarsuaq plain south of bulletin 14: gsb191-indhold 04/12/07 14:37 side 37 38 ilimanaq (claushavn) and ilulissat (figs 22, 25; table 3). this area can be characterised as an upland with altitudes up to 400–600 m a.s.l.; the mountain ridges are partly tillcovered and in the intervening valleys, marine silt is overlain by glacio-fluvial deposits related to the subsequent tasiussaq moraine system. a straightforward correlation of the numerous moraine remnants cannot be made. outwash deposits and beach ridges up to 60–70 m a.s.l. related to the most westerly of these moraine remnants can now be dated to around 10–9 ka b.p. according to the emergence curves of the area (long et al. 1999). the deposition of the basal silt must also be related to this period. the three dated deposits south of kangia of 9.9 (ua-4574) and 9.8 (k-2023 and k-1818) ka b.p. are all situated in west-facing coves (fig. 25). the northernmost dated sample from near ilulissat of 9.3 ka b.p. (ua-1086) was taken from marine silt underlying the archaeological site at sermermiut (fig. 26). the archaeology and palaeobotany of this site have been described by larsen & meldgaard (1958) and fredskild (1967). farther north, there is clear evidence that the western coast of arveprinsen ejland (alluttoq) was ice-free before 9.7–9.9 ka b.p. (long et al. 1999), which implies that the position of the marrait moraine system could be at the mouth of the torsukattak fjord (fig. 22). the younger tasiussaq moraine system in disko bugt can be followed from the nuuk peninsula, mentioned above and described by long & roberts (2002), northwards along most of the bay. morphologically, it is characterised by marginal moraines and wide alluvial plains. the dated sites are from south to north (fig. 25): 1. the kangersuneq fjord, 22 km north of the nuuk peninsula, where a shell sample gave a date of 7.9 ka b.p. (ua4572, table 3). the shell material was collected in a cliff 15–20 m a.s.l. near serfarsuit at the head of the kanger suneq fjord. the date gives a minimum age of the deglaciation of this site. fig. 24. upland landscape at c. 600 m a.s.l. between kangia and paakitsoq fjord, east of central disko bugt, looking towards the north-west. the person is standing on the central part of an interlobate moraine (‘the fjord stage’). the moraine is a part of a system of ice-margin features extending from the iceberg bank at the mouth of kangia north-eastwards to the mouth of paakitsoq. the present margin of the inland ice is visible in the right background. this area was investigated for a potential hydro-power plant in the 1980s. r.e. peary and c. majgaard visited the ice margin here in 1886. photograph by a. weidick in 1963. bulletin 14: gsb191-indhold 04/12/07 14:37 side 38 39 2. two samples from a locality 15 km farther to the north in eqaluit (laksebugt) gave dates of 8.7 and 8.5 ka b.p. (ua-4573 and k-993, table 3; fig. 25). both are from marine silt covered by gravel of an alluvial plain at altitudes of 38–40 m a.s.l. according to recent detailed mapping of the area by geus. the elevation of the sampling sites, determined by altimeter, was stated to be 50–55 m by weidick (1968). the alluvial plain is related to the ice margin features of the tasiussaq moraine system. 3. from south of claushavn (ilimanaq) and north of ilulissat, dates on marine shells have been obtained by different authors (table 3; fig. 25). in addition to the group of dates between 9.9 and 9.3 ka b.p. mentioned above, a younger group of shell samples, dated to 8.6–7.7 ka b.p., appear to be related to a marine level around 40 m. the marine sediments at this elevation were laiddown before the alluvial plains of the tasiussaq moraine system. this is most markedly seen at the narsarsuaq (‘lersletten’) plain (table 3), where details of braided rivers and dead ice holes have been recognised (weidick 1968). two exceptions to the younger dates should be noted: (1) a basal gyttja from a lake sediment core in an oxbow lake of the alluvial plain of narsarsuaq gave an age of 8.8 ka b.p. (k-987; kelly in: tauber 1968), which is older than the underlying marine sediments (k-992, 7.9 ka b.p.). the date of the lake sediment is thus presumed to be too old, possibly due to hard-water effects or reworked older carbon. (2) a shell sample dated to 7.7 ka b.p. that was collected at an elevation of 30 m a.s.l. (i-6243, table 3); this sample is presumed to relate to a marine level at or below this height. near qajaa, a minimum age for the deglaciation is provided by gyttja dated to 8.8–8.0 ka b.p., and the tasiussaq moraine system must be older than this. farther north, there is evidence that paakitsoq was ice-free before 7.7 ka b.p. (fig. 22; long et al. 2006). the marine limit at these sites was found to be around 40 m a.s.l. little is known about the age of the moraine systems in the northern parts of disko bugt. two radiocarbon dates of marine shells from the inner parts of the fjords in this region are available. one was obtained from the coast near the southern flank of the outlet eqip sermia (fig. 22), 51°w 10 km kangia narsarsuaq sermermiut inland ice eqaluit paakitsoq ua-4577: 3.1 ua-699: 3.2 ua-4576: 3.4 ua-1085: 4.6 ua-1086: 9.3 k-2023: 9.8 ua-4574: 9.9 i-6243: 7.7 ua-4573: 8.7 k-993: 8.5 k-2022: 8.6 k-992: 7.9 k-1818: 9.8 k-987: 8.8 ua-4575: 8.7 ua-4572: 7.9 tasiusaq qajaa ilimanaq serfarsuit paakitsoq tissarissoq ua-4581: 3.4 ua-4580: 3.8 ua-4582: 3.8 ua-4583: 4.0 ua-2350: 4.3 ua-4579: 5.6 ua-4578: 6.1 alanngorliup sermia ua-1087: 2.2 ua-4585: 2.6 ua-1088: 3.9 ua-4584: 5.2 50°w 69°30´n 69°n fig. 25. radiocarbon dates related to holocene ice-margin deposits in the kangia region. the older group (dates in bold: 9.9–9.3 ka b.p.) is related to the marrait moraine system whereas younger dates (8.7–7.7 ka b.p.) are related to the tasiussaq moraine system. the 6.1–2.2 ka b.p. dates along the ice margin relate to recession during the holocene thermal maximum. for details, see tables 3 and 4. note that the map is based on 1994 icemargin data (cf. fig. 13). bulletin 14: gsb191-indhold 04/12/07 14:37 side 39 40 where shells collected 2–3 m a.s.l. yielded an age of 7.3 ka b.p. (k-6373; rasch 1997), while another shell collection from qapiarfiit on the south side of the outlet kangilerngata sermia (fig. 22) at the same elevation gave an age of 8.4 ka b.p. (k-3663; ingólfsson et al. 1990). it appears, therefore, that the attainment of the present ice-margin position in the fjords in this region here was reached as early as about 8 ka b.p. this implies that the zone of moraine deposits in this region diverges, so that the northern correlatives of the tasiussaq moraine system are found in the interior parts of the fjords, close to the present ice margin. a high concentration of ice-contact features is usually interpreted as indicative of deposition at a ‘stable’ ice margin. the dates of the ‘fjord stage’ cover a time of c. 2 ka (c. 9.9–7.9 ka b.p.), with a net recession of the ice margin in the central and southern parts of disko bugt of only a few kilometres. the two-fold division of the fjord stage into the older marrait moraine system and a younger tasiussaq moraine system might reflect two phases of development, with the older primarily due to decreased ablation and calfice production following the reduced contact between the sea and the ice margin. another factor is the complexity of the response of the ice margin to climate change. the response of a continental ice sheet and its marginal positions depend on long-term changes in ice flow (on centennial to millennial time scales) due to sustained changes in accumulation and surface temperatures. however, short-term annual to decadal changes in mass-balance elements, such as accumulation, run-off and iceberg calving, also play important roles (reeh 1999; dahl-jensen 2000). all these elements are present in the case of the early holocene change of the ice cover in disko bugt. thus the abrupt increase in snow accumulation over greenland at the end of the younger dryas, as documented by alley et al. (1993), may have had a positive effect on the mass balance and response of the ice margin, which could counter the effect of the subsequent temperature increase. the abrupt temperature rise of 10–15ºc from around 11.7 to c. 10 ka b.p., that caused the break-up of the ice cover over disko bugt at c. 10 ka b.p., must have led to a change of the ice-margin profile. increased ablation, especially at lower levels of the ice, would lead to a steeper slope 1 km sermermiut glacial striae marine silt silt-rich till moraine boulders boulder ridge kangia 51°07´w 69°12´n roads in ilulissat l ll ll l l l l l fig. 26. sketch-map of the area around ilulissat (jakobshavn) and sermermiut showing quaternary features. the map is based on the geodetic institute (now part of national survey and cadastre, copenhagen) map sheets 1:2000, jakobshavn and 1:8000, sermermiut. contour interval is 10 m. redrafted from weidick (1969). bulletin 14: gsb191-indhold 04/12/07 14:37 side 40 41 of the ice margin and also to a reduction of the ablation zone, until changes in the ice dynamics led to a new quasiequilibrium. studies of the effects of the present centennial temperature rise on the ice margin indicate that such changes of geometry are currently taking place (thomas et al. 2001; bøggild et al. 2004; hughes 2004). during the build-up of both the marrait and the tasiussaq moraine systems (c. 9.9–7.9 ka b.p.), the front of jakobshavn isbræ was resting on or at isfjeldsbanken at the mouth of kangia. iceberg production, estimated from the frontal area determined by the depth of the iceberg bank (200–300 m) and the trend of the moraines, must have been reduced during this period (weidick 1994a, b). during the subsequent recession, the ice margin may have reached the position of the present location at or before 6–7 ka b.p. palaeoceanographic investigations show that large volumes of meltwater produced by jakobshavn isbræ deflected the warmer waters of the west greenland current away from the eastern coastal areas of disko bugt until 7.9 ka b.p., when the glacier finally receded from isfjeldsbanken to the interior of the present kangia fjord system. this recession allowed the warmer west greenland current to penetrate to the eastern part of disko bugt, as can be seen from changes in the fauna (lloyd et al. 2005). the relationship between the recession of the ice margin and the warming of the bay must be rather complex. after the recession from the tasiussaq moraine system, where numerous drainage channels led to the formation of alluvial plains with the main drainage from the ice-sheet margin directly to disko bugt, the drainage became concentrated into a few channels that drained into the deep fjords of kangia and torsukattak. the first slow-down of the ice recession is marked by the marrait moraine system (around 9.9–9.4 ka b.p.), which probably developed in response to the change of environ51˚w 50˚w 69˚30´n 69˚n ilulissat qasigiannguit 9500 9500 9500 8000 a.d. 2000 a.d. 1850 8000 8000 5000–4000 5000–4000 land area deglaciated after the little ice age (150 years b.p.) fjord deglaciated after the little ice age land area glaciated after 5000–4000 years b.p.ice margin at maximum recession (5000–4000 b.p.) marrait (m) and tasiussaq (t) moraine systems 10 km m t m t fig. 27. provisional reconstructions of the position of the ice margin in the ilulissat area at c. 9500, 8000, 5000–4000 and 150 years b.p. the change of the ice margin between the little ice age maximum (150 years b.p.) and the present day (the trimline zone) is shown in red for land areas and horizontal shading for floating glaciers. note that the zone without vegetation (the trimline zone) becomes narrow south of jakobshavn isbræ and almost absent farther south. bulletin 14: gsb191-indhold 04/12/07 14:37 side 41 42 ment and a much reduced output of calf ice in the central and southern parts of disko bugt, and possibly also a changed profile of the ice margin. the halt or slow-down may have been prolonged during the deposition of the tasiussaq moraine system (around 8.8–7.9 ka b.p.), influenced by the 8.2 ka b.p. cold event (fig. 23; o’brien et al. 1995; dansgaard 2004). this event was the most marked cold episode during the early holocene, and in greenland it was probably associated with the formation of the moraine on the nuuk peninsula in the south-east corner of disko bugt (long & roberts 2002). the 8.2 ka b.p. event was probably related to the release of large amounts of cold meltwater from the receding laurentide ice sheet. the subsequent recession of the ice-sheet margin from the tasiussaq moraine system to its present position is not known in detail. the recession began at around 8 ka b.p., and the attainment of the present ice-margin position is estimated to have occurred at or before 6–7 ka b.p. for the jakobshavn isbræ area (weidick et al. 1990). this is based on the minimum age of the tasiussaq moraine system of about 8 ka b.p. (fig. 27), and on dates of marine material (mainly shells) transported westwards by the ice to the present margin of the ice; the age range of these dates is shown in fig. 25. it can be seen that the dated samples cover a time span from 6.1 to 2.2 ka b.p., during which time the ice margin was east of its present position. when attempting to correlate the individual stages in the recession and the subsequent neoglacial advance of the ice margin, the scant evidence from the regions north and south of disko bugt should also be considered. in the uummannaq fjord complex to the north, it is known that the ice margin was situated well into the fjord complex at about 10.7–10.5 ka b.p. (símonarson 1981; bennike 2000; bennike & björck 2002). the high relief of the fjord landscape, the deep fjords, and the many still productive calving glaciers (fig. 5) might well explain an early recession, although the time at which the glaciers reached their present position is unknown. south of disko bugt, the lowlands east of aasiaat provide little information about the recession history of the ice margin. the southern shores of disko bugt were deglaciated prior to 10.5 ka b.p. (donner & jungner 1975; long et al. 2003). a moraine system is found c. 25 km west of the present front of nordenskiöld gletscher (fig. 2), and seems to be related to the large alluvial plain of naternaq (lersletten). the alluvial plain overlies marine deposits (e.g. harder et al. 1949; laursen 1950), but little systematic work has been carried out here. the altitudes of the lakes on this plain are 37–54 m a.s.l. (geus 2004), and it is presumed that the tasiussaq moraine system relates to a sea level about 50–40 m above the present day. a morphological extension of these moraines to the south might correspond to the locality near ussuit (67º51´n, 50º16´w; fig. 1), from where kelly (1973) obtained a date of 7.6 ka b.p. on marine shells. these shells appear to be related to a marine level of c. 42 m a.s.l., and provide a minimum age for the adjacent moraine system. as the onset of deglaciation of the outer coast south of disko bugt took place well before c. 10.5 ka b.p., the subsequent net recession of the ice margin probably occurred at approximately the same rate as in the northern parts of disko bugt. however, local temporal variations in the recession must have characterised the region. during the recession over the lowlands between arfersiorfik fjord in the south and disko bugt in the north, the relative sea level was 50–100 m above the present. although an extended contact between the ice margin and the sea can be envisaged, the region was a shallow area compared to the depths in disko bugt. in the area south of disko bugt, attainment of the present ice-margin position has been estimated to have occurred at about 8 ka b.p., with the ice margin subsequently receding eastwards beyond its present position (letréguilly et al. 1991b). the calculation for the holocene thermal maximum at around 5 ka b.p. gives a recession of the ice margin in this region of the same order of magnitude as that recorded at jakobshavn isbræ. the reason for the high sensitivity of the ice margin to climate change may be that this region is characterised by extensive coastal areas of low elevation and low accumulation. the position of the ice margin is thus essentially controlled by ablation, resulting in a rapid response to any early holocene warming. in fjords or bays (such as disko bugt), as well as in lowland areas such as those south of disko bugt, the rapid holocene recession is mainly related to increased ablation. other factors, as noted above, may be important, and can explain local temporary deviations from the general case; these must be considered when constructing a more detailed history of ice-margin changes. roberts & long (2005) have suggested a convergent iceflow drainage pattern through the shallow troughs in the southern part of disko bugt, and through the shallow fjords south of disko bugt, related to a large drainage outlet through the outer egedesminde dyb. it is not clear, however, when ice drainage of this magnitude could have taken place. the ice-sheet margin after the holocene thermal maximum collections of marine fossils from little ice age moraines at, and close to, the present ice margin have been made bulletin 14: gsb191-indhold 04/12/07 14:37 side 42 43 around jakobshavn isbræ. the detailed subsurface maps of this area make it possible to follow the trends of the fjords beneath the ice, to establish the glacial transport route of the dated material, and hence to calculate that at maximum recession, the ice margin was some 15–20 km eastwards of the present position (weidick et al. 1990; weidick 1992a). the ages of 6.1–2.2 ka b.p. provide a minimum estimate for the period during which the ice margin at most localities was situated east of its present position (table 4; fig. 25). the ice margin presumably gradually advanced after the end of the holocene thermal maximum at 5–4 ka b.p., although exceptions may have occurred locally. the narssarssuaq moraine system near narsarsuaq in south greenland (weidick et al. 2004b; bennike & sparrenbom 2007) and the drygalski moraines (fig. 22) crossing the root of the nuussuaq peninsula north of disko bugt (fig. 22; kelly 1980), may be related to early neoglacial events. if so, the margin of the ice sheet may locally have advanced beyond the present position in some places during early phases of the neoglacial. the age of the narssarssuaq moraine system is estimated to be c. 2 ka b.p. (weidick et al. 2004b; bennike & sparrenbom 2007), whereas the age of the drygalski moraines is unknown. investigations of sediment cores sampled west of isfjeldsbanken indicate maximum atlantic water influence during the period from c. 1650 to 500 calendar years b.p., which is related to a recession of the front of jakobshavn isbræ during the medieval warm period (lloyd 2006). modelling of ice-margin changes for the last 1400 years is based on information on the surface and subsurface topography of the ice margin, palaeoclimate data from ice cores, measured weather data, and the rheology of the ice (fig. 28). the calculations by reeh (1983) demonstrate that quite short lengths of the ice margin exhibit local variations in behaviour. two major advances took place, one at about a.d. 800 and another from a.d. 1500 to 1900 related to the little ice age. the latter was often associated with advances, notably at about 1750 and in the late 1800s. both of these advances can be observed in the response curves, with the older advance often apparently the more prominent. however, older moraines are rarely observable in the field, since they are often buried beneath younger ones. the difference between the c. 1750 and the c. 1900 maxima is only about 100–200 m. for the last 100–200 years, the modelled ice-margin changes agree with observational data, because fresh deglaciated terrain with a width of about 1 km is found. this 1–2 km width of the ‘historical advance and recession’ of the ice after the middle of the 19th century is commonly quoted, but in reality large variations occur. an unusually large width is seen around jakobshavn isbræ (fig. 29), where fresh moraines and ice-polished bedrock locality/material laboratory 14c age† (± 2σ) rcorr14c age§ calib. age¢ δ13c reference number* years b.p. years b.p. ka b.p. ‰ paakitsoq, c. 69°25´n, 50°20´w macoma calcarea shell ua-4577 3300 ± 65 2900 3.1 –6.79 this study shell ua-699 3420 ± 105 3020 3.2 0¶ weidick et al. (1990) mytilus edulis shell ua-4576 3560 ± 65 3160 3.4 –3.73 this study shell ua-1085 4520 ± 135 4120 4.6 0¶ weidick et al. (1990) tissarissoq, c. 69°06´n, 50°02´w mya truncata shell ua-4581 3590 ± 65 3190 3.4 1.88 this study hiatella arctica shell ua-4582 3940 ± 65 3540 3.8 1.92 this study mya truncata shell ua-4580 3945 ± 70 3545 3.8 2.50 this study mya truncata shell ua-4583 4075 ± 70 3675 4.0 2.16 this study odobenus rosmarus tusk ua-2350 4290 ± 100 3890 4.3 –13.05¶ weidick (1992a) mya truncata shell ua-4579 5240 ± 75 4840 5.6 1.85 this study balanus sp. plate ua-4578 5710 ± 55 5310 6.1 0.98 this study alanngorliup sermia, c. 68°54´n, 50°15´w shell ua-1087 2620 ± 110 2220 2.2 0¶ weidick et al. (1990) mya truncata shell ua-4585 2935 ± 60 2535 2.6 2.01 this study shell ua-1088 4000 ± 115 3600 3.9 0¶ weidick et al. (1990) mya truncata shell ua-4584 4930 ± 60 4530 5.2 1.78 this study table 4. neoglacial radiocarbon age determinations (see also table 2, appendix 2) * ua: ångström laboratory, uppsala. † the radiocarbon age determinations have been corrected for measured or assumed isotopic composition by normalising to –25‰ on the pdb scale. § rcorr: reservoir-corrected. corrected for a seawater reservoir effect of 400 years. ¢ calibrated using the intcal04 dataset (reimer et al. 2004) and the oxcal. v.3.10 software program (bronk ramsey 2001). ¶ assumed value. bulletin 14: gsb191-indhold 04/12/07 14:37 side 43 44 (the trimline zone) extend for over 30 km in front of the present glacier. the thinning of the ice, estimated from the height of the trimline zone, is 200–300 m around jakobshavn isbræ (figs 27, 29). by contrast, just c. 25 km south of jakobshavn isbræ, the trimline zone nearly disappears around the outlets of alanngorliup sermia and saqqarliup sermia (fig. 30). historical records indicate nearly stationary conditions here since the middle of the 19th century (weidick 1994a, b); the glaciers almost maintain their maximum extent from the little ice age. further south, towards kangerlussuaq at 67ºn, many lowland sectors were characterised by a minor readvance in the period between c. 1950 and 1985 (weidick 1992a, 1994a, b); this could be a consequence of the temperature fall in the last decades of the 1900s, or that the ice sheet had still not adjusted to past climatic fluctuations, as modelled by huybrechts (1994). the recent development of the trimline zone around jakobshavn isbræ has been studied from multispectral landsat images (csatho et al. 2005). observations of glacier change are unevenly distributed, with maximum change in south-west greenland that appears to be indicative of a present warming trend. this is confirmed by observations and measurements in south greenland (between 60º and 65ºn; mayer et al. 2002; podlech 2004), while investigations in melville bugt (73º to 79ºn, c.e. bøggild, personal communication 2004) show the same warming trend; a similar situation may apply to jakobshavn isbræ (see also below). in general, the little ice age has left its mark in the form of fresh moraines, which are well defined in some regions, but not in others. exact dating of parts of these moraines clearly points to a specific geological event; the link between moraine formation and climate, however, is not always straightforward. for a regional understanding of ice-margin history, only the whole zone or belt of icemargin deposits can serve as a useful comparison with models of the ice-sheet response to climate change. 49°w49°30´w 50°w 69°30´n 69°45´n 20 km 800 advance recession 1000 1200 1400 1600 1800 1 0 –1d is ta nc e (k m ) years (a.d.) 1400 1100 1200 1300800 700600 900 1000 disko bugt 50 km sector boundary (see inset map) sub-sector boundary n fig. 28. advances and recessions in the ice-sheet margin can be correlated with climatic variation deduced from the dye 3 ice-core record. lateral variation in the response was defined for a specific sector of the margin, north of jakobshavn isbræ (see inset map); sub-sectors (indicated by colours) show varying degrees of response, indicated on the inset, lower right. from reeh (1983). bulletin 14: gsb191-indhold 04/12/07 14:37 side 44 45 a b c fig. 29. a: front of jakobshavn isbræ (northern side). the vegetation-poor zone (the trimline zone) is 200–300 m high. photograph by j. lautrup 1991. b, c: south side of jakobshavn isbræ; highest mountain is 368 m. the photograph in b (by m.c. engell) is from 1902 whereas that in c (by a. weidick) is from 1963. in 60 years, the glacier front has receded about 13 km to the east, and is seen faintly in the distance on the 1963 photograph. bulletin 14: gsb191-indhold 04/12/07 14:37 side 45 46 it is not known to what extent the fluctuations of the ice margin in the paakitsoq area can be correlated with the fluctuations of jakobshavn isbræ. however, during the period from 4.6 to 3.1 ka b.p. (fig. 25), the ice margin at paakitsoq was situated east of its present position. at c. 1200 years b.p., the ice margin advanced and reached a position near the subsequent little ice age position. it is supposed that a gradual net advance of the ice margin took place after the thermal maximum that ended around 4 ka b.p. although this was interrupted by minor stillstands or recessions, according to the modelling of reeh (1983). evidence for late holocene fluctuations has also been recorded from eastern disko bugt (lloyd 2006). the beginning of the little ice age advance may be related to the legend about tissarissoq, the ice-filled bay south of kangia. the name tissarissoq is claimed to refer to a time when hunting in the bay was possible, that is to say when glacier ice did not cover the bay (hammer 1883). this accords with a gradual glacier advance during the 1700s. however, of the 15 shell samples so far dated from the ice margin around jakobshavn isbræ (fig. 25; table 4), the youngest radiocarbon age (2.2 ka b.p.; ua-1087) is a millennium before the thule culture arrived in the region. relative sea-level changes around disko bugt thule winter houses and norse ruins were reported to be partly below sea level by thorhallesen (1776) and arctander (1793), indicating recent submergence. these observations and further investigations in the first half of the 19th century by pingel (1841, 1845) showed that this submergence had been preceded by emergence, as indicated by raised marine deposits. for the area around disko bugt, systematic descriptions of former raised shorelines and measurements of the following submergence were initiated in the last half of the 19th century (steenstrup 1883a, b; saxov 1958). from the 1950s onwards, detailed investigations of changes in sea level, mapping of the marine limit and descriptions of marine faunas were carried out by ggu (laursen 1950; donner & jungner 1975; weidick 1975, 1976). this work has been followed up in the past few decades with more comprehensive studies of relative sealevel changes, often with support from the arctic station in godhavn (ingólfsson et al. 1990; bennike et al. 1994; rasch & jensen 1997; long et al. 1999, 2003, 2006; rasch fig. 30. western margin of alanngorliup sermia in 1988. the glacier has a steep margin close to vegetated terrain. photograph by a. weidick. (this glacier is recorded as avannarleq sermeq by johansen & nielsen 2001.) bulletin 14: gsb191-indhold 04/12/07 14:37 side 46 2000; long & roberts 2002, 2003). much of the accumulated data was reviewed by fleming (2000), and used for comparisons with results of geophysical modelling of the ice-sheet history in greenland (tarasov & peltier 2002; fleming & lambeck 2004). the relative sea-level changes observed in greenland are mainly related to the combined effects of local glacio-isostatic responses of the earth’s crust to variations in glacier load, and global eustatic changes of sea level due to the storage and melting of ice on the continents. in the disko bugt region, the holocene recession was complete by c. 6–5 ka b.p., and was followed by a neoglacial expansion of the ice cover. the other major ice sheets in the northern hemisphere disappeared, the fenno-scandinavian ice sheet at 10–9 ka b.p., and the laurentide ice sheet at about 8–7 ka b.p. in the antarctic, holocene recession of the shelf ice that began at the end of the last ice age has continued until the present day (bindschadler & bentley 2002); although, modelling predicts expansion during the next few centuries (huybrechts et al. 2004). although c. 40% of the volume of the greenland ice sheet has disappeared since the lgm (huybrechts 2002), a substantial glacier load is still present in the central part of greenland. the main losses of the glacier load after the lgm have occurred at the present ice margin (fig. 31). by contrast, where other ice sheets have completely disappeared, the maximum glacio-isostatic uplift has occurred in what was formerly their central part. the change of relative sea level has been dominated by holocene emergence caused by recession and thinning of the ice margin. the altitude of the marine limit, which is defined as the maximum height of relative sea level after the last deglaciation, is usually determined by the upper limit of raised shorelines and/or the lower limit of perched boulders, or by studies of sediments in lakes situated above and below the marine limit. the trend of the marine limit indicates an elongated dome over the outer ice-free land, parallel to the present coast, with the highest values for the altitude of the marine limit in areas that show the largest holocene recession of the ice-sheet margin (fig. 32). higher values for the altitude of the marine limit than indicated on fig. 32 were reported in some parts of the disko bugt region by rasch (2000) and long et al. (2006). these 47 el ev at io n (m ) bedrock 54° 52° 50° 3000 2000 1000 0 –1000 48° 46° 44° longitude (°w) greatest change in ice load disko bugt 4000 9000–8000 6000 and present 10 000 fig. 31. provisional simplified profile of icemargin stages between 10 000 years b.p. and the present day from south of disko island to jakobshavn isbræ (sermeq kujalleq). the approximate position where the greatest change in glacier load has taken place is indicated. from weidick (1993). 500 km 120 120 80 80 120 40 40 20 40 80 80 20 40 80 fig. 32. elevation of the marine limit in greenland in metres a.s.l. compiled from weidick (1992b), bennike & weidick (2001), bennike (2002) and weidick et al. (2004b). bulletin 14: gsb191-indhold 04/12/07 14:37 side 47 48 workers defined the marine limit largely on the basis of the lower limit of perched boulders. it should be acknowledged, however, that the shores in the region may locally have been affected by large waves so that the present lower limit of perched boulders may not be a true reflection of sea level. large waves can be generated at glacier fronts by calving of icebergs, or by turnover of icebergs. the so-called kanelling, which occurred in the harbour of ilulissat in the early parts of the 20th century (reeh & engelund 1971; reeh 1985), was characterised by far-travelled large waves. such waves have been described from several localities in greenland; they are typically recorded from narrow fjords with calving outlets, and can reach heights over 10 m (reeh 1985). in addition, landslides have generated tsunamis in areas with steep slopes around the vaigat strait and the eastern shores of disko. in 2000, a tsunami that resulted from a landslide in the vaigait strait had a run-up height of 50 m close to the landslide and a run-up height of 28 m at a distance of 20–25 km from the slide (dahl-jensen et al. 2004). minimum values for the marine limit in the central parts of disko bugt can be deduced from the uppermost marine terraces and beach ridges, which are found at 70–80 m a.s.l. this corresponds to an age of around 10–9 ka b.p. according to the relative sea-level curves of long et al. (2006), which is close to the minimum ages for the last deglaciation (fig. 22). while the number of observations has increased considerably in recent years, the previously recorded maximum value of the marine limit south of disko bugt has been largely confirmed. however, the different versions of the marine limit that have appeared over the past decades are still only generalised views, based on an uneven distribution of observations (rasch 2000, long et al. 2006). data coverage is relatively good around disko and disko bugt, but very scattered farther north. only few observations are therefore available for the uummannaq fjord complex. future detailed mapping of the marine limit will probably give a more varied picture (e.g. ingólfsson et al. 1990). a fall in the marine limit from 85 m to 54 m a.s.l. over a distance of 8 km was reported for the south-eastern corner of disko bugt by long & roberts (2002), the change being attributed to a slowdown of ice-margin recession (cf. the fjorde stage). the dating of the limit at any one locality is usually based on extrapolation of local relative sea-level curves. the age of the marine limit in west greenland decreases from the outer coastland towards the east (fig. 33). westwards, an apparent convergence of strandlines is seen, although their trend is often uncertain. a reverse trend has been suggested for the west coast of disko (funder & hansen 1996), such that the marine limit becomes younger westwards. this was based on the occurrence of transgressive sequences in the area (ingólfsson et al. 1990), and is comparable to transgressions that have been reported from western norway (andersen 1965; kaland et al. 1984). the form and trend of the marine limit are determined locally by the former glacier load and by the subsequent rate of thinning and recession of the ice cover. detailed determinations of local sea-level changes, and the spatial trends of the individual isobases, are important for understanding the development of the landscape. compared to other regions in greenland, the amount of detailed data on relative sea-level changes in the disko 54°w distance from outer coast (km) present inland ice margin 20050 150 100 250 50 0 1000 4 ka b.p. 5 6 ka b.p. 7 8 910 marine limit inland ice margin c. 4 ka b.p. el ev at io n (m a .s .l. ) 1 el ev at io n of ic esh ee t su rf ac e (k m ) 2 50 km 70°n 69°n 52°w 50°w fig. 33. hypothetical shoreline relation diagram for disko bugt broadly following the 69°n line of latitude (see inset map); the position of the margin of the greenland ice sheet at the present day and at c. 4 ka b.p. is indicated. note that the dashed shorelines for 4 ka and 8 ka b.p. indicate the probable trends of these shorelines prior to the neoglacial ice-sheet advance and accompanying proximal depression of the shoreline. bulletin 14: gsb191-indhold 04/12/07 14:37 side 48 49 bugt area is large (rasch 1997, 2000). the older relative sea-level curves were mainly based on dates of marine shells (donner & jungner 1975; donner 1978; weidick 1996; rasch & jensen 1997). the oldest dates at any locality provide minimum ages for the local deglaciation and of the marine limit. however, the relationship between localities with fossil marine shells and the contemporaneous sea level is somewhat uncertain, and numerous sample localities from a large area are needed to provide enough data points. a more recent series of detailed curves has been constructed from isolation basins (long et al. 1999, 2003, 2006; long & roberts 2002, 2003). the constructed uplift curves indicate a steady emergence throughout the early and midholocene, followed by a late holocene submergence, presumably caused by the advancing ice margin and increasing glacier load. the relative sea-level curves indicate a larger initial emergence to the east, near the present inland ice margin, than farther west, and the hypothetical shoreline diagram has been drawn on this basis (fig. 33). the north–south trend of the marine limit suggests the trend of the isobases should be broadly parallel to the present ice-sheet margin, but locally a more complicated pattern than shown in fig. 33 can be expected (rasch 2000). a more exact and site-specific method of dating relative uplift is by dating of the timing of isolation of lakes at different altitudes (fig. 34). this procedure has been carried out at six localities around disko bugt (fig. 2): the vaskebugt (kangerluarsuk) area on arveprinsen ejland (long et al. 1999), akulliit/nuuk in the south-east corner of disko bugt (long & roberts 2002), qeqertarsuatsiaq in southwestern disko bugt (long & roberts 2003), innaarsuit on southern disko (long et al. 2003), and near qajaa and at paakitsoq in eastern disko bugt (long et al. 2006). the main drawbacks of this method are that it is time consuming, and obviously it can only be applied to areas where lakes exist at different elevations below the marine limit. regional correlation of locally determined uplift may well be substantiated through geomorphological correlation of strandlines in the area. from the investigations referred to above, it has been established that the present sea level was reached by 5–4 ka b.p. emergence continued for some time, and the lowest relative sea level was reached at around 3–1 ka b.p. when it was about 5 m below the present; this was followed by the beginning of the present submergence. the period after c. 4.5 ka b.p. coincides with the period of human settlement in greenland, and a number of the earliest known ruins are at or below water level at present high tide (larsen & meldgaard 1958; rasch & jensen 1997). the exact form of the late holocene part of the relative sealevel curves is difficult to establish because the curves are flat, and the transition from emergence to submergence took place at shallow depth. at tuapaat on southern disko island, morpho-stratigraphic investigations of the coastal landscape by rasch & nielsen (1995) suggested that three or four transgressions have taken place during the past 2.5 ka b.p. sea-level measurements were initiated at godhavn in 1897, and demonstrate a subsidence of 0.475 m up to 1946, whereas an emergence of 0.3 m was recorded for the time period between 1946 and 1957 (saxov 1958; kelly 1980). on the basis of repeated gps observations in west greenland between 1995 and 2002, the present-day vertical crustal movements have been determined (dietrich et al. 2005). at ilulissat an uplift of 1.6 mm/year was observed, which is presumably due to the recent thinning of the inland ice in this region. uplift was also recorded at the outer coast south of disko bugt. in contrast, marked subsidence characterised the inland region south of disko bugt, with rates up to 4 mm/year. 10 20 0 50 60 30 40 70 80 90 100 10 8 6 4 2 0 age (cal. ka b.p.) a lti tu de ( m a .s .l. ) present sea level south-western disko bugt north-eastern disko bugt 54°w 50 km disko bugt 70°n 69°n 52°w 50°w fig. 34. examples of two relative sea-level curves from the disko bugt area. the inset map shows the location of the sites in the disko bugt region. modified from long et al. (1999) and long & roberts (2003). bulletin 14: gsb191-indhold 04/12/07 14:37 side 49 50 from the geological section, it is clear that a presentation of the history of the jakobshavn isbræ ice stream must also include the history of the surrounding parts of the ice sheet. in the same way, the present major ice streams are a part of the local glaciation history of disko bugt. subsurface of the ice-sheet margin the first mapping of large parts of the ice sheet was made in the period 1949–1951 during the ‘expéditions polaires françaises’ using seismic methods (holtzscherer & bauer 1954). their surveys covered the southern parts of the inland ice (south of c. 70ºn in west greenland and south of c. 72º40´n in east greenland) and indicated a depression below present sea level in the central part of the ice sheet as well as a drainage channel towards disko bugt (figs 4, 17). later airborne radar surveys between 1968 and 1976 provided a measure of the thickness of the entire ice sheet (gudmandsen & jakobsen 1976; overgaard 1981). these data provided a realistic impression of the gross features of the subglacial terrain, including highlands, uplands, lowlands, and drainage areas (figs 4, 17). more recent data on the subsurface of the ice sheet are steadily improving our knowledge of the subglacial landscape (bamber et al. 2001). in general, the elevation of the subglacial terrain falls from the marginal areas towards the central depression of the ice sheet, and as noted above, a depression connects kangia with the interior region, lying at or below sea level, (fig. 4). ice streams have been defined as “part of an ice sheet, in which the ice flows more rapidly and not necessarily in the same direction as the surrounding ice” (armstrong et al. 1973, p. 26). the strong flow of ice streams, relative to the slower moving ice on either side, leads to strong and chaotic break-up of the ice in ice streams. this limits pene tration by radar waves, and hence the determination of ice thickness below ice streams. however, the thickness of jakobshavn isbræ has been determined by seismic methods (clarke & echelmeyer 1996). in the central parts of this ice stream, the ice thickness varies from c. 1.9 km near the grounding zone (cf. fig. 35) with the ice surface at an altitude of c. 500 m a.s.l., to about 2.5 km at a distance of 40 km behind the grounding zone where the ice surface is c. 1000 m a.s.l. thus the bottom of the subglacial trough is found at a depth of c. 1.5 km below sea level; the trough can be described as a canyon-like feature about 7 km wide, surrounded by a hilly subglacial landscape with elevations close to sea level. much of the basal interface is probably underlain by compacted, non-deformable sediments (clarke & echelmeyer 1996). farther inland, the trough beneath the ice stream gradually levels out. about 120 km east of the grounding zone, the subsurface depression can be interpreted as a shallow subglacial valley. the subsurface map of fig. 4 also shows that other subglacial areas near the present ice margin in west greenland are dominated by uplands, with no clear indication of other ice streams of the size of jakobshavn isbræ. supplementary depth soundings have provided more detailed information around the deep drilling sites in the central parts of the ice sheet and in marginal areas studied in connection with possible exploitation of hydropower resources. for example, a c. 800 km2 area of the ice margin in the paakitsoq area, north of ilulissat, was mapped at a scale of 1:250 000 with 100 m contour intervals (thomsen et al. 1988). other less detailed maps cover smaller areas of the ice margin at tissarissoq (south of ilulissat), alanngorliup sermia and saqqarliup sermia (thorning et al. 1986; thorning & hansen 1987; weidick et al. 1990). one revelation from this detailed mapping was that even for ice thicknesses of 600–800 m, the depressions and elevations of the subsurface are reflected in the topography of the ice surface, although in a smoothed and somewhat distorted form. it thus appears that topographical features of the ice-free marginal areas bordering the ice sheet continue beneath the ice, and are readily discernable on landsat images with a low sun angle (fig. 36). the deep trough of jakobshavn isbræ is clearly visible from its abundant crevasses (dark colour, relative high ablation), in contrast to the other less important ice streams draining into disko bugt (sermeq kujalleq and sermeq avannarleq in torsukattak icefjord). the continuation of torsukattak icefjord beneath the ice sheet rapidly levels off into the subglacial uplands of the area, as seen from the depth of the present fjord (c. 600 m) and radar soundings 25–30 km from the front (overgaard 1981); this picture is confirmed by the image of fig. 4. restricted ablation and high relief may explain the local high production of calf ice from the glaciers draining into uummannaq fjord farther to the north (fig. 5). snowfall here is also heavier than in the interior parts of disko bugt (ohmura & reeh 1991). glaciology bulletin 14: gsb191-indhold 04/12/07 14:37 side 50 51 present ice-margin surface a topographic map of greenland at a scale of 1:2 500 000 with 250 m contour intervals, including the entire ice sheet, was published by kms [national survey and cadastre] in 1994, and the same topographic base was used in the geological map of greenland at the same scale (escher & pulvertaft 1995). these maps show a low surface gradient from the highest central part of the ice sheet outwards towards the ice margin, reflecting in some respects the lowland areas underlying the ice (fig. 4). detailed surface maps were produced in connection with the investigations along the epf-egig line at eqip sermia (holtzscherer & bauer 1954), over the ice sheet margin at paakitsoq (thomsen et al. 1988), and around jakobshavn isbræ (fastook & hughes 1994; fastook et al. 1995). the map around jakobshavn isbræ was used to delineate the ‘ilulissat ice d ep th /a lti tu de ( km ) distance (km) distance (km) distance (km) distance (km) sea level ice surface s n s n s n s n 202 202 4 202 4 202 4 2 0 –2 2 0 –2 2 0 –2 2 0 –2 b 1 2 3 10 km 51°w 69°30´n 69°15´n 69°n 4 1000 1000 1500 1500 500 500 500 1000 1500 1 2 3 4 kangia ilulissat oqaatsut ilimanaq a 50°w 49°w 10 km boundary of ilulissat icefjord world heritage site ice-stream grounding zone ice-stream boundary fig. 35. a: map of kangia and the ablation zone of jakobshavn isbræ showing the positions of cross-sections 1–4 illustrated below. the ice-free areas have been mapped by geus, and the inland ice by h. brecher (in: fastook & hughes 1994 and fastook et al. 1995); supplementary data from the inland ice have been provided by s. ekholm (national survey and cadastre, copenhagen). map compilation by willy weng and annette t. hindø (geus). the borders of the jakobshavn ice stream (black lines) and the grounding zone c. 1985 (dotted black line) are from echelmeyer et al. (1991, 1992). ice-margin data from 1994; note that the present margin of jakobshavn isbræ, following the dramatic recent recession (fig. 13), broadly coincides with the grounding line shown here. b: the four profiles of the ice stream (1 to 4) are based on seismic studies (clarke & echelmeyer 1996). n and s on the profiles indicate the north and south ends. bulletin 14: gsb191-indhold 04/12/07 14:37 side 51 fjord’ world heritage proposal that was included on the world heritage list in 2004 (mikkelsen & ingerslev 2002). this map is reproduced here (fig. 35), and clearly shows the trend of jakobshavn isbræ. in the disko bugt region, annual mass-balance field measurements have been carried out along the epf-egig line (holtzscherer & bauer 1954; bauer et al. 1968b, ambach 1977), the ggu stake line (fig. 15; thomsen et al. 1988; braithwaite et al. 1992) and a line along jakobshavn isbræ (figs 37, 38; echelmeyer et al. 1992). the altitude of the equilibrium line, where the annual mass balance is 0, shows great annual variations, from c. 1000 m a.s.l. (extremely cold budget years) to more than 1200 m a.s.l. an in-depth discussion of climatic factors determining the variations in annual mass balance is given by echelmeyer et al. (1992), who also record the albedo changes (increased ablation) due to inblown dust from the extensive trimline zone around jakobshavn isbræ. calculated variations in 52 50°w 49°w 48°w 69 °n 47°w 70 °n 69 °3 0´ n 51°w 50°w 49°w 70 °n 69 °3 0´ n 52°w 70°30´n 70 °3 0´ n 48 °w 40 km fig. 36. false-colour satellite image (landsat-7) of the interior of disko bugt and adjacent parts of the ice sheet, from 4 march 2002. the outer part of kangia, and coastal areas north and south of the mouth of kangia, are ice-free (black areas in lower part of the image). the morphology of the snow-covered, ice-free land continues under the surface of the ice margin, as can be seen from the surface features of the ice. bulletin 14: gsb191-indhold 04/12/07 14:37 side 52 53 the annual ablation of the region for the period from 1961/1962 to 1989/1990 are reported by braithwaite et al. (1992). the trend of the annual mass balance along the ggu line, measured in 1982/83 and 1983/84 (thomsen et al. 1988), is compared in fig. 37 with the mass balance along jakobshavn isbræ, measured from 1984 to 1988 (echelmeyer et al. 1992). the trends of the curves (the ablation gradient) are similar, but the difference between the equilibrium line altitudes (ela) mainly illustrates the variations of the annual climatic conditions in the area, rather than variations due to the distance between the locations of the measurements (fig. 38). the ela defines the upper limit of the ablation area, which is about 50 km wide. above this, the accumulation zone extends to the top of the ice sheet around summit. the decrease of intermittent summer-snow melt with increasing elevation is expressed by the division of the firn area into a region of superimposed ice over wet snow, a percolation facies and finally a dry snow facies. these facies were originally defined by benson (1962), and have been modified by williams et al. (1991) and benson (1994, 1996). the facies concept is applied in the current monitoring programme of the mass balance of the inland ice, which is based on data from satellites. the spectral variability of ice and snow surfaces is used to determine the different facies (fausto et al. 2007). the meltwater produced above the ela during the summer is retained in the firn in increasing amounts with decreasing elevation. refreezing of this meltwater leads to 70°n peary and majgaard 1886 epf egig a. w eg en er ex pe dit ion eg ig ggu de quervain 1912 1200 1200 1200 600 600 600 800 800 800 1000 1000 1000 1400 1400 1400 1600 1600 1600 1800 1800 1800 ilulissat eqip sermia 25 km disko bugt 69°30´n 51°w 50°w 49°w 48°w fig. 38. map of the jakobshavn isbræ area, showing routes, stations and stake lines of the expeditions that have worked on the ice. egig, expédition glaciologique internationale au groenlande, mainly 1957–1960; epf, expéditions polaires francaises (missions p.e. victor 1948–1951); ggu, geological survey of greenland, stake line 1982–1989. modified from wegener et al. (1930) and weidick & thomsen (1983). el ev at io n (m a .s .l. ) 900 1000 1200 1300 1400 1100 800 700 600 500 400 300 200 –4 –3 –2 –1 0 1 2 1500 water equivalent (m) 19 83 /8 4 19 82 /8 3 19 84 /8 8 fig. 37. annual mass balance in relation to elevation along profiles in the ilulissat area. 1982/83 and 1983/84 designate profiles along the ggu stake line (see figs 15 and 38 for location; from thomsen et al. 1988). 1984/88 shows measurements for this time period along jakobshavn isbræ (from echelmeyer et al. 1992). bulletin 14: gsb191-indhold 04/12/07 14:37 side 53 54 increased temperatures in the firn. in the ablation zone, the superficial meltwater drains through crevasses and subglacial tunnel systems, where refreezing and closure of the meltwater conduits can take place. the complex system of drainage that can occur in glaciers has been described by roethlisberger & lang (1987). descriptions of the drainage in the paakitsoq area (site of the ggu line) are provided by thomsen et al. (1988) and zwally et al. (2002), and for jakobshavn isbræ by echelmeyer et al. (1992). drainage and thermal conditions at the ice margin the large ablation area of a continental ice sheet such as the inland ice is characterised by extensive englacial and subglacial drainage. superficial meltwater drains into subglacial channels and conduits that are fed from crevasses and moulins on the ice surface. at the base of the ice, the water can drain in thin water films or in subglacial channels to the ice margin (thomsen et al. 1988), and the water can emerge as upwelling plumes at the front of calving tidewater glaciers (outlets) such as jakobshavn isbræ (echelmeyer et al. 1992). the contribution of basal melting from the jakobshavn isbræ drainage basin has been calculated to be c. 20% of the total loss by ablation (echelmeyer et al. 1992). the complexity of the drainage is influenced by refreezing or by internal as well as external heating due to the movement of the glacier. glacier-dynamic modelling has been applied to the socalled ‘quiet’ sector, located between jakobshavn isbræ in the south and the ice streams draining into torsukattak icefjord in the north. the ice margin in this sector is bordered by land areas or lakes. modelling of the bottom conditions by radok et al. (1982) suggested that the basal ice reaches its pressure melting point 250–300 km from the ice margin. based on a model for calculating the response of the marginal sector of the inland ice to mass-balance changes, the ice margin at paakitsoq (fig. 28) could be divided into three zones: (1) an inner zone more than 292 km from the ice margin where ice movement is dominated by internal deformation. (2) a zone between 292 and 18 km from the margin where the ice is at the pressure-melting point, which leads to a significant degree of bottom sliding. (3) an outer zone, at 18–0 km from the ice margin, that is characterised by extensive sliding, high hydraulic pressure at the bottom of the ice, and conditions that allow for formation of cavities (reeh 1983; thomsen et al. 1988). ice-temperature measurements have been made at a few localities along both the epf-egig and ggu lines. at camp vi on the epf-egig line, drilling from the surface at 1598 m a.s.l. extended only to a depth of 125 m; a temperature profile from this hole shows a decrease in temperature from c. –12.5ºc to –16ºc (heuberger 1954; robin 1983). the ice thickness here was about 1000 m. on the ggu line, temperatures were measured in a hole drilled at c. 500 m a.s.l. temperatures were slightly negative throughout the ice hole, decreasing to –2.1ºc in the upper 50 m (thomsen et al. 1991). for jakobshavn isbræ, it has been calculated that 2–3 km3/year of meltwater is generated by deformational heating, although the meltwater does not seem to influence the movement of the ice stream (echelmeyer & harrison 1990; echelmeyer et al. 1992). d ep th ( m b el ow s ur fa ce ) sub-glacial surface 1200 1000 600 400 200 800 1400 1600 1800 2000 2200 2400 2600 –20–25 –15 –10 –5 0 temperature (°c) fig. 39. temperature profile through jakobshavn isbræ near its centre line, showing temperature versus depth below the ice surface (from iken et al. 1993). the profile is situated about 50 km from the glacier front of 1988/89, at an elevation of 1000 m a.s.l., with the bottom of the ice at about 1500 m b.s.l. the dashed lines show possible limits for temperatures below the measured values. more recent modelling indicates that the thickness of the basal temperate ice (ice at the pressure-melting point) is around 230 m in the ice stream, and c. 30 m in the adjacent ice sheet (funk et al. 1994). bulletin 14: gsb191-indhold 04/12/07 14:37 side 54 55 a temperature profile of jakobshavn isbræ was meas ured in 1988/1989 at a locality situated about 50 km from the glacier front at 1020 m a.s.l. the site is located at the centre line of the ice stream, where the surface ice moves at about 1 km/year, and the bed is situated at about 1500 m below sea level (iken et al. 1993, clarke & echelmeyer 1996). the temperature profile (fig. 39) shows a tempera ture minimum of –22ºc c. 1200 m below the surface; this implies a thick, relatively warm and low viscosity bottom layer which is thought to facilitate the fast movement of the ice stream. it has been suggested that this basal layer may contain wisconsinan ice (iken et al. 1993). this work was followed up by investigations of flow and temperature conditions at the transition between jakobshavn isbræ and the ‘quiet’ ice margin (lüthi et al. 2002). the surface features of the ‘quiet’ sectors of ice comprise lakes, rivers, crevasse formations and moulins, which have been mapped in detail and described along the lower parts of the egig line by bauer et al. (1968a). detailed maps and descriptions around the ggu line near paakitsoq are 50°w 49°w 48°w 69 °n 47°w 70 °n 69 °3 0´ n 51°w 50°w 49°w 70 °n 69 °3 0´ n 52°w 70°30´n 70 °3 0´ n 48 °w 40 km fig. 40. false-colour satellite image (landsat-7) of the interior of disko bugt and adjacent parts of the ice sheet, from 7 july 2001. the scene shows the zone with numerous lakes on the ice near the equilibrium line. bulletin 14: gsb191-indhold 04/12/07 14:37 side 55 56 given by thomsen et al. (1988). thorough descriptions of jakobshavn isbræ and its surroundings are provided by echelmeyer et al. (1992) and fastook et al. (1995). the surface of the inland ice margin in the disko bugt region is characterised by numerous lakes up to an altitude of about 1400 m (fig. 40). above this altitude, corresponding approximately to the beginning of the wet snow facies of the accumulation area, lakes are present but are ice and snow covered and only faintly visible. the occurrence of numerous lakes extends down to about 1000 m a.s.l., corresponding to the zone of superimposed ice and the upper part of the ablation area. lower down in the ablation area there are fewer lakes, due to the decreasing thickness of the ice cover that leads to increased crevassing during ice movement. where the ice margin is bordered by land, marginal lakes are common. due to the bordering hilly upland, the occurrence and size of the lakes vary with changes in the position of the ice margin, just as the proglacial drainage shows variations. changes in drainage patterns at paakitsoq, between 1953 and 1959, can be documented by aerial photographs from the 1950s, 1960s and 1985 (thomsen 1983). tininnilik, situated c. 40 km south of jakobshavn isbræ, is an ice-dammed lake at the ice margin that shows periodic drainage. the lake covers an area of 43 km2 at its maximum extent and 20 km2 at its minimum (thomsen 1984). when the periodic drainage was first described in 1913 (koch & wegener 1930), it was stated that the filling/drainage cycle of the ice-dammed lake was about 10 years based on information from local people. this cycle seems to be more-or-less permanent, probably due to the nature of the damming glacier (saqqarliup sermia) which, in contrast to jakobshavn isbræ, has been almost stable over the last 150 years (weidick 1994a, b). braithwaite & thomsen (1984) determined that during each drainage event nearly 2 km3 of water is released along saqqarliup sermia into the southern branch of the tasiusaq fjord complex, and from there onwards to the south side of kangia. braithwaite & thomsen (1984) also recorded the years of drainage for the period 1942–1983; the most recent drainage took place in 2003 (f. nielsen, personal communication 2004). 3 0 100 200 300 400 500 2 1 0 ice-sheet surface particle path (flow line) ice divide accumulation zone equilibrium line ablation zone summit distance (km) a lti tu de ( km ) younger dryas younger dryas bølling 1500 –40–45 –35 –30–40 –35 –30 δ18 o (‰)δ18 o (‰) 2000 2500 3000 denekamp hengelo glinde oerel odderade brørup eemian saalian? holsteinian? 35 ka b.p. 60 ka b.p. 100 0 200 300 400 500 600 700 ba b a d is ta nc e (m ) a ge ( ka b .p .) d ep th ( m ) 10 11 12 14 16 18 20 25 30 35 40 50 60 70 80 90 100 150 200 250 fig. 41. cross-section of the inland ice, showing schematic flow lines within the ice and the locations of (a) an oxygen isotope record near the margin and (b) the deep ice core at the summit of the ice sheet. a: profile of the outer ice margin at paakitsoq. the record covers major parts of the last ice age (wisconsinan). from reeh et al. (1991). b: a section of the ice core at summit (fig. 1) revealing successively older layers of snow and ice. the record covers about 250 000 years, and includes the eemian and perhaps holsteinian interglacials (dansgaard et al. 1993). bulletin 14: gsb191-indhold 04/12/07 14:37 side 56 57 the age of the ice margin has been investigated in detail along three profiles near paakitsoq (fig. 2). at certain localities along the ice margin, the ice stratification demonstrated from deep drill holes in the interior of the ice sheet can be preserved (fig. 41; reeh et al. 2002, petrenko et al. 2006). such localities, as at paakitsoq, are characterised by a fairly smooth subsurface within so-called quiet marginal areas distant from larger ice streams. the profiles so far investigated at paakitsoq cover a span of time from possibly 150 000 years b.p. to the present. these and similar investigations at the ice margin provide easy access to ice-age ice, a cheaply acquired supplement to the very expensive deep ice-core records, which can provide important information about past climates and the dynamics of the ice sheet. movement of the ice margin the surface movement of the ice margin in disko bugt has been determined in the ‘quiet’ land-based area along the egig line north of jakobshavn isbræ, and in fig. 42 it is compared with the horizontal movement of jakobshavn isbræ. the main differences are found in the areas around and below the equilibrium line. the thickness of the landbased ice is up to 700–800 m below the ablation zone along the egig profile (holtzscherer & bauer 1954; bauer et al. 1968b). along this profile, the horizontal surface movement increases westwards to a maximum of 100–200 m/year just above the equilibrium line, and then decreases towards the ice margin. at the swiss station (for location, see fig. 15), measurements of movements were made from 1996 to 1999 (zwally et al. 2002). the station is situated near the equilibrium line at 1175 m a.s.l., at a site where the ice is 1220 m thick. the measurements show increasing ice velocities with increasing surface melt, indicating that bottom sliding is enhanced by rapid migration of meltwater to the bottom of the ice. this provides a mechanism for rapid, large-scale dynamic response of ice sheets to climate warming, for example during the transition from glacial to initial interglacial conditions. with respect to calving tidewater outlets in fjords as well as calving outlets in proglacial lakes, it is generally known that the rate of movement increases towards the front reaching up to a few km/year. however, little information is available for calving outlets in the disko bugt region apart from at jakobshavn isbræ. 5000 ve lo ci ty ( km /y ea r) 7 6 5 4 3 2 1 0 distance from margin (km) 3 2 1 0 –1 el ev at io n (k m ) 3 2 1 0 –1 el ev at io n (k m ) ? ve lo ci ty ( km /y ea r) distance from margin (km) 0 500 0 equilibrium line ice sheet ice sheet 1 a b fig. 42. a: diagram showing a crosssection through jakobshavn isbræ together with a plot of the horizontal surface movement of the ice stream. compiled from fastook et al. (1995) and joughin et al. (2004). b: cross-section of the ice sheet along the egig line (fig. 38) in a ‘quiet’ marginal area, about 60 km north of jakobshavn isbræ, in association with a plot of the horizontal surface movement of this part of the ice sheet, which has a maximum velocity of 200 m/year c. 100 km from the ice margin (from bauer et al. 1968b). bulletin 14: gsb191-indhold 04/12/07 14:37 side 57 58 jakobshavn isbræ is an extreme example, where the depth of the glacier is controlled by a subglacial trough reaching 80–100 km inland under the ice and with depths that reach down to 1.5 km below sea level in the outer parts (clarke & echelmeyer 1996). the trough can be envisaged as a continuation of the proglacial icefjord (kangia), which is believed to be around 1000 m deep. such a depth is extreme for the fjords in the area, where depths of 400–800 m are more usual (weidick et al. 1974b). fjord depths at the present fronts are not known, and can only be estimated from soundings made at some distance from the active fronts. the empirical relationship between calving rate and water depth at the glacier front (pelto & warren 1991) is therefore difficult to establish, although an attempt was made by long & roberts (2003) for the deglaciation of disko bugt. measurements of movement and thickness of all the tidewater glaciers in disko bugt and the uumannaaq fjord complex were undertaken in 1957 and 1964 by the egig expeditions (bauer et al. 1968a; carbonnell & bauer 1968), based on photogrammetric analyses of aerial photographs, and show a positive correlation between mean rate of movement and mean thickness of the glacier front. alanngorliup sermia, july 1875 (helland 1876) <0.5 <0.2? jakobshavn isbræ, 7–9 july 1875 (helland 1876) from south side 0.40 14.7 5.4 0.42 15.4 5.6 0.45 15.2 5.5 0.45 15.2 5.5 1.05 19.8 7.2 1.06 19.5 7.1 jakobshavn isbræ, 22 march – 24 april 1880 (hammer 1883) from south side 0.28 5.1 1.9 0.55 7.5 2.7 0.62 9.2 3.4 0.88 12.5 4.6 0.87 12.3 4.5 jakobshavn isbræ, july 1902 (engell 1904) from south side 1.29 15.0 5.5 1.30 14.2 5.2 1.87 19.8 7.2 1.84 19.8 7.2 4.26 22.8 8.3 jakobshavn isbræ, 27–28 september 1929 (sorge in wegener et al. 1930) from south side 2.1–3.7 18–21 6.6–7.6 6.6–7.7 sermeq avannarleq in torsukattaq fjord, 24–25 july 1875 (helland 1876) from north side 0.21 3.8 1.4 0.37 5.7 2.1 1.93 8.8 3.2 4.07 10.1 3.7 4.94 10.2 3.7 4.97 9.4 3.4 sermeq avannarleq in torsukattaq fjord (steenstrup 1883a) from north side, 2.70 7.8 2.8 5–7 may 1879 2.70 6.3 2.3 from north side, 3.01 5.0 1.8 21–22 may 1880 3.01 7.8 2.8 2.17 5.0 1.8 1.51 5.0 1.8 distance from flank of glacier (km) velocity comments m/24 h km/y table 5. velocity measurements of glaciers in the disko bugt region before 1950 velocity measurement uncertain measurements close to front frontal height scarcely over 40 m a.s.l. width of the fjord estimated to 4.5 km (on present maps c. 7 km) measured c. 5 km behind glacier front maximum velocity in the central part of the glacier estimated to at least 16 m/24 h. height of glacier front c. 63 m a.s.l. width of fjord c. 7 km velocity measured near the front width of fjord c. 7 km preliminary calculations based on four points. glacier c. 6 km wide, presumably measured near front presumably measured near front recorded frontal height 15 m a.s.l. glacier c. 9 km wide measured near glacier front width of glacier estimated to c. 8 km bulletin 14: gsb191-indhold 04/12/07 14:37 side 58 59 with respect to the high movement rate of many of the calving glaciers in the disko bugt region, and variations of the velocity over time, a brief historical review is given below. the field conditions for the earliest measurements introduce an element of uncertainty, but do provide an order of magnitude of the possible variations in velocity over long time-spans. all the measurements given are related to the frontal areas, and for nearly all the glacier fronts the variations in position are within 2 km for the period since the end of the 19th century (fig. 43; weidick 1994a, b). the measurements of velocity at the front of most outlets were thus conducted at nearly the same location. the only exception is jakobshavn isbræ, which receded 26 km between c. 1850 and c. 1950 and has experienced continued recession since 2002/2003; in contrast to other outlets, the measurements were here undertaken from widely different positions of the glacier front during the recession. most early velocity measurements were carried out over short time intervals, and consequently where a low rate of movement was recorded, possible errors may be large, a fact that is stressed in some of the old descriptions (helland 1876). in table 5, older data are only given for the faster moving glaciers. some original sources gave measurements in danish feet (1 danish foot = 0.31385 m), here converted to metres or kilometres. rates of movement are given as m/24 h or km/year (tables 1, 5). the latter is used for comparison with modern data and neglects possible variations in the course of the year. notes on individual outlets the glacial histories of individual glacier outlets in the disko bugt region (figs 2, 43) are summarised below. nordenskiöld gletscher (akuliarutsip sermersua); 68°20´n, 50°51´w). this glacier does not drain into disko bugt, but its glacial history is closely related to the bay. the visits of the nordenskiöld expeditions to the glacier in 1870 and 1883 gave rise to detailed descriptions of this glacier (nordenskiöld 1885, 1886), but the velocity of the outlet was apparently first measured in july 1957 (bauer et al. 1968a). the average velocity was found to be 3 m/24 h with a small production of calf ice (fig. 5). from the descriptions, it appears that the build-up of a frontal moraine and proglacial delta hinders the production of calf ice. in the period 2000–2005, a slight increase in velocity was reported by rignot & kanagaratnam (2006). saqqarliup sermia (68°54´n, 50°18´w) and alanngorliup sermia (68°55´n, 50°12´w). alanngorliup sermia was visited in 1875 by helland (1876), who found that the front had a height of c. 10 m. the velocity of the glacier was given as below 0.5 m/24 h, probably recorded in the lower reaches of the glacier although the location was not stated. higher up in the glacier, where there is a tributary to saqqarliup sermia, the rate of movement is given as a maximum of 0.4 m/24 h. for both outlets, the mean frontal velocity in 1957 was measured at 0.9 m/24 h (carbonnell & bauer 1968). a slight velocity increase has also been reported for this outlet by rignot & kanagaratnam (2006). jakobshavn isbræ (sermeq kujalleq; 69°11´n, 49°48´w). the first velocity measurements in 1875 were made along the southern side of the glacier. in the fast flowing, central part of the glacier, c. 1 km from the margin, a velocity of 19.8 m/24 h was measured (table 5; helland 1876). helland also measured the movement of the glacier adjacent to the margin, where he found a velocity of not more than 0.02 m/24 h. subsequent measurements in 1880 recorded slightly lower velocities (table 5; hammer 1883; estimated maximum velocity >16 m/24 hours). the measurements appear to have been made c. 5 km behind the front, judging from a sketch map of measured points in hammer’s report. if the up-stream velocity decrease was similar to later values (carbonnell & bauer 1968), the frontal velocity may well have been of the same magnitude as given by helland. however, the velocity does not always decrease immediately behind the glacier front; joughin et al. (2004) reported that the marked decrease in speed of jakobshavn isbræ in the 1990s first occurred c. 14 km behind the front. in 1902, the glacier was visited by engell, who recorded a similar high velocity to that given by helland for the central part of the glacier (c. 23 m/24 h; engell 1904). comparable values of between 18 m/24 h and 21 m/24 h were measured by sorge in 1929 (wegener 1930; wegener et al. 1930). measurements in july 1958 gave a mean velocity of 13.1 m/24 h (bauer et al. 1968a), whereas investigations in june 1964 gave a mean velocity of 19.1 m/24 h (carbonnell & bauer 1968). a very similar figure to the ‘frontal velocities’ of the central zone of the glacier given by pelto et al. (1989): 21.1 (1964), 20.4 (1976), 21.0 (1978) 20.6 (1985) and 20.3 (1986) m/24 h. a decrease of velocity immediately behind the front was documented. in a comment to the surprisingly large difference between the velocities in 1958 and 1964, it was pointed out that similar large changes were found at rink isbræ (fig 1.), and that more measurements are needed to explain the difference (carbonell & bauer 1968, p. 77). subsequent measurements of jakobshavn isbræ show velocity variations bulletin 14: gsb191-indhold 04/12/07 14:37 side 59 60 lille gletscher store gletscher sermeq kujalleq sermeq avannarleq sermeq avannarleq kangilerngata sermia eqip sermia alanngorliup sermia saqqarliup sermia nordenskiöld gletscher jakobshavn isbræ usulluup sermia a.d. 1850 2 km 2 km 2 km 3 km 14 16 18 20 22 24 26 km 3 km 4 km 3 km 2 4 6 8 10 12 2 km 3 km 2 km 3 km 1900 1950 2000a.d. 1850 1900 1950 2000 inngia isbræ umiammakku isbræ rink isbræ kangerlussuup sermersua kangerluarsuup sermia sermeq avannarleq perlerfiup sermia sermiq silarleq kangilleq sermilik upernavik isstrøm 4 km 2 4 6 4 km 3 km 8 km 4 km 3 km 4 km 3 km 3 km 2 km 22 km 20 18 16 14 12 10 8 6 4 fig. 43. tentative reconstructions of fluctuations in the frontal positions of calving glaciers in central west greenland from 1850 to 1985. upernavik isstrøm shows a recession of c. 23 km and jakobshavn isbræ shows a recession of c. 26 km (red lines), whereas the other glaciers show smaller fluctuations. from weidick (1994b). bulletin 14: gsb191-indhold 04/12/07 14:37 side 60 61 that are apparently related to the thickness of the ice margin. however, information is needed on the velocity in the time period between the measurements in 1929 and in 1948. it is also possible that the low velocity in 1958 could have been connected to a thickening of the glacier from c. 1950–2000, which was a period with a stable front. studies of aerial photographs from the 1940s may provide data on thickness changes in the marginal parts of the ice sheet. it is noteworthy that the maximum velocity, in spite of all the possible errors that may have affected the older measurements, has maintained a value of c. 5–9 km/year for more than a century. this is particularly relevant when considering the reasons for the marked velocity increase of the glacier after about 2000. thus a 95% velocity increase took place between 1996 and 2005 according to rignot & kanagaratnam (2006). more details are provided below and in fig. 44. in 2003, the velocity was 12.6 km/year and the discharge was c. 50 km3 ice/year (joughin et al. 2004). sermeq avannarleq in kangia (69°2´n, 50°18´w). the velocity of this glacier was first measured during the egig expeditions, who recorded an average velocity of c. 1 m/ 24 h (carbonnell & bauer 1968). eqip sermia (69°48´n, 50°13´w). velocity measurements of this glacier were first made from 31 july to 7 august 1912 by de quervain, who recorded a maximum velocity at the front of 1.5–2.4 m/24 h (de quervain 1925). subsequent more detailed investigations on velocity and frontal fluctuations were made by bauer during the epf expeditions in 1948–1949 (bauer 1955). a mean velocity of 3 m/ 24 h was recorded, and no difference was noted between september 1948 and june 1949. the egig measurements from 12 to 17 july 1957 gave a mean velocity value of 3.1 m/24 h (bauer et al. 1968a), whereas later measurements, from 5 to 9 july 1959 (bauer 1968) and 29 june to 12 july 1964 (carbonnell & bauer 1968) both gave values of c. 2 m/24 h. the differences are related to insufficient measuring points for the old data (bauer 1968, p. 10). detailed records of the frontal fluctuations of the glacier in the 20th century are provided by bauer (1955) and nielsen et al. (2000), who indicate advances around 1920, and during the 1990s. between 2000 and 2005, the velocity of eqip sermia accelerated by 30% (rignot & kanagaratnam 2006). kangilerngata sermia (69°55´n, 50°17´w). the movement of this glacier was first measured from 7 to 17 july 1957, when the average velocity was given as 2.3 m/24 h (bauer et al. 1968a). subsequent measurements from 9 to 22 june 1964 gave 3.3 m/24 h (carbonnell & bauer 1968). between 2000 and 2005, the velocity of kangilerngata sermia also increased by 30% (rignot & kanagaratnam 2006). sermeq kujalleq in the torsukattak icefjord (70°00´n, 50°19´w). velocity measurements of this large glacier started with the measurements of bauer et al. (1968a) from 12 to 17 july 1957, which indicated a mean velocity of 7.2 m/24 h. from 9 to 22 june 1964, a mean velocity of 9.7 m/24 h was recorded (carbonnell & bauer 1968). sermeq kujalleq slowed down by 11% between 2000 and 2005 (rignot & kanagaratnam 2006). sermeq avannarleq in the torsukattak icefjord (70°04´n, 50°19´w). the velocity was first measured by helland (1876) from 24 to 25 july 1875. the measurements were made from the northern side of the glacier, and extended 4 km into the central parts where a velocity of over 10 m/24 h was measured (table 5). subsequent measurements by steenstrup (1883a) in may 1879 and may 1880, from nearly the same position, did not reach as far into the gla cier, but a velocity of 8 m/24 h was recorded c. 3 km from the glacier margin. this can be compared with later measurements of the mean glacier velocity from 12 to 17 july 1957 of 6.4 m/24 h (bauer et al. 1968a) and on 9 to 22 june 1964 of 5.2 m/24 h (carbonnell & bauer 1968). the velocity profile of the glacier is irregular. maximum vel ocities of 9.5 m/24 h was found c. 1.5 km from the gla cier margin (carbonnell & bauer 1968, fig. 38). sermeq avannarleq, as sermeq kujalleq, slowed down between 2000 and 2005 by 11% (rignot & kanagaratnam 2006). sermeq kujalleq (store gletscher: 70°24´n, 50°32´w). although situated somewhat farther to the north, north of disko bugt, this glacier (‘store qarajaq bræ’ of steenstrup 1883a and ‘grossen qarajaq eisstrom’ of von drygalski 1897) deserves mention because of its history of explor ation. its velocity was measured in august 1878. the maximum velocity of 12 m/24 h, measured c. 3 km from its north side by steenstrup (1883a), is close to the mean velocity recorded by von drygalski (1897) in 1893 of 12.9 m/24 h, of 11.6 m/24 h by bauer et al. (1968a) for 1957, and by carbonnell & bauer (1968) of 13.4 m/24 h for 1964. these authors noted a significant upstream decrease in velocity for this glacier. historical information on the measurements of the calving tidewater glaciers in and around disko bugt leaves the general impression of high-speed behaviour of the major outlets to the fjords in the region; the scattered investigations lack sufficient details for further conclusions. the bulletin 14: gsb191-indhold 04/12/07 14:37 side 61 62 early measurements, especially those of von drygalski’s ‘grossen qarajaq eisstrom’ and other outlets to disko bugt, reported the significant upstream decrease of the glacier velocities, although the investigations only covered the outermost 5–10 km of the glacier lobe. jakobshavn isbræ is characterised by a high and rather constant flow rate at the glacier front, even though the position of the front has changed with time. detailed coverage of the areal distribution of velocity is only known for jakobshavn isbræ, in the form of velocity contours (isotachytes of ahlmann 1948). these cover an area of c. 80 × 80 km of the ice sheet upstream of the glacier front, with contours of 50 m/year, and record the situation at around 1985 (fastook et al. 1995). the profile of fig. 42 is derived from this source. a revision of the map for february 1992 and october 2000 (i.e. before the collapse of the front of jakobshavn isbræ in 2002/2003), is provided by joughin et al. (2004), who describe the subsequent development illustrated by movement profiles from the front and reaching 50 km inland. remote sensing data provide details of the frontal changes from 1985 to the present. the velocity of jakobshavn isbræ has been more variable during the past few years than at any time since records began. a net thinning of the glacier of 200–300 m over the past 150 years, corresponding to an average thinning of 1.3 –2.0 m/year, is apparent from the elevation of the trimline zone around the glacier. this thinning has taken place over a time period when the glacier has had a rather constant velocity of 5 to 9 km/year. detailed records covering the past few decades indicate that the 1984 velocity of 6.7 km/year had decreased to 5.7 km/year by 1992, and that this lower velocity continued until 1997. the vel ocity then increased sharply, and for the years 2000, 2002 and 2003 the glacier attained velocities of respectively 9.4, 11.9 and 12.6 km/year (joughin et al. 2004). a thickening of c. 1 m/year was related to the period of low speed (1991–1997), whereas the subsequent higher velocities were related to a thinning of around 6 m/year, fig. 44. the 15 10 5 1985 1992 1994 1995 2003200220012000 ve lo ci ty ( km /y ea r) velocity (km /year) yea r (a .d.) 10 20 30 40 10 20 30 40 10 20 30 40 50 1990 year (a .d.) 1995 2000 1985 15 10 5 10 20 30 distance (km from front) 40 50 thicke ning thinning fig. 44. variations in the rate of movement of jakobshavn isbræ at distances c. 5–50 km behind the front. the marked increase in velocity that began in the late 1990s coincided with thinning of the front. simplified from joughin et al. (2004). bulletin 14: gsb191-indhold 04/12/07 14:37 side 62 63 a b 5 km fig. 45. frontal area of jakobshavn isbræ. a: landsat image of 27 september 1979. the glacier front and released icebergs can be seen near the left margin of the image. the ice stream is characterised by linear structures. a tributary from the north is separated from the main stream by a subglacial rumple. south of the front of jakobshavn isbræ is an area of stagnant ice (tissarissoq, see also figs 2, 25) that is separated from the ice stream by another rumple. the arrow indicates the view of the photograph below. b: photograph of jakobshavn isbræ viewed from the wsw (see arrow on fig. 45a) where the northern tributary joins the main ice stream. the bedrock topography under the ice is clearly reflected in the topography of the ice surface. photograph by h.h. thomsen, 1984. bulletin 14: gsb191-indhold 04/12/07 14:37 side 63 64 sudden transition to rapid thinning that followed was at first confined to areas below c. 500 m a.s.l., but then spread inland and by about the year 2000 had reached up to 2000 m a.s.l. (thomas et al. 2003). the apparent quasi-stability of the front of jakobshavn isbræ from c. 1950 to 2000 has been related by echelmeyer et al. (1991) to the presence of pinning points in the frontal areas of the floating front. however, little is known about the depths of the fjord below the floating outer parts of the glacier front that lie c. 22 km west of the lower seismic station of clarke & echelmeyer (1996), corresponding to profile 1 in fig. 35. a possible threshold near the grounding zone (figs 36, 45) may be viewed as a northern continuation of the curved bedrock lineaments on the south side of the glacier. in contrast to the dramatic changes of the thickness and position of jakobshavn isbræ, the other calf-ice producing outlets to the north and south have shown only small changes in frontal positions during the past 150 years (weidick 1994a, b). these small changes are perhaps linked to the shallower depths of these outlets demonstrated by the radar surveys of the technical university of denmark (overgaard 1981) and apparent from low sun angle landsat scenes. for the areas around jakobshavn isbræ, the ice margin seems to have been nearly continuously receding over the period from c. 1850 to 1950. marginal zones farther to the north and south of jakobshavn isbræ, however, show a slight advance during the last decades of the 20th century (fig. 43). the regional monitoring of glaciers in west greenland on the basis of aerial photographs stopped with the last full coverage flown in 1985; subsequent regional coverage has been based on satellite imagery. modelling of present response patterns of the ice sheet to climatic changes seems to match the present patterns of recession and readvance, and an important thickening of the south-western parts of the inland ice seems to have taken place (huybrechts 1994). in contrast to the modelling, monitoring of outlet glaciers and marginal elevation changes based on repeated surveys by laser altimetry in 1993/1994 and 1998/1999 has revealed a significant thinning of the surroundings of jakobshavn isbræ, whereas the thickness of jakobshavn isbræ itself was constant or increasing (abdalati et al. 2001). however, it may be misleading to compare the very generalised trends of the change of the ice margin positions based on scattered historical information with the detailed trends revealed during the past few decades (see e.g. thomas et al. 2003 and joughin et al. 2004 for jakobshavn isbræ). with respect to the response of ice streams from the inland ice to climatic forcing, a holistic modelling approach has been presented by hughes (2004), which lengthens and lowers the profiles of the greenland ice streams. bulletin 14: gsb191-indhold 04/12/07 14:37 side 64 65 the account of the onset and subsequent, repeated glaciations of greenland, and the disko bugt region in particu lar, leaves more questions than answers. however, the detailed investigations of recent decades support the old idea that the preglacial, major fluvial drainage pattern of central greenland was westwards towards the disko bugt region. the subsequent drainage of the greenland ice sheet since its formation and during repeated glaciations was also predominantly westwards, and at the maximum extent of the inland ice, major ice streams in the present offshore region occupied and modified former river valleys and fjords. the onset and extent of the early glaciations are still not clear. they must, however, have had characteristics similar to the glaciation of the present antarctic or to high-arctic ice shelves, in that they showed a high sensitivity to climatic and eustatic sea-level changes and perhaps included ice streams of different character from those that drain the inland ice today. during the illinoian, the glacial limit of the ice sheet may have been located at store hellefiskebanke and at the central part of disko banke. during the wisconsinan, the ice margin may have extended only to the eastern part of store hellefiskebanke. on both occasions, the outer egedesminde dyb was probably occupied by a high-arctic type ‘ice stream’, but data to support this scenario are so far lacking. a number of interglacial and interstadial deposits have been discovered along the outer coast of disko island, and farther north at other localities along the extreme western parts along the outer coast. these have been referred to a number of marine events, but the extent of the ice sheet during and between the events is unknown. modelling indicates, however, that during the last interglacial, the eemian or sangamonian, the inland ice was reduced to a degree where it was almost separated into a southern and northern ice sheet (fig. 14). the extent of the ice sheet during the last glacial maximum at 21 ka b.p. is still not established, neither from offshore stratigraphy nor from dates from marine sediment cores. it is likely that the pronounced warming at 14.7 ka b.p., combined with an initial rise of global sea level, caused a recession of the ice margin. during the cold interval of the younger dryas, the ice margin may have been located at the marked basalt escarpment that forms a submarine barrier, which would mean that the marine ice margin was situated at a depth of only 300–400 m. in spite of the depth of this barrier, it must have had a blocking effect on ice movement, so that the large piedmont glacier that filled disko bugt during and immediately after the younger dryas was thinning rather than receding. in contrast, the few radiocarbon dates from vaigat, north of disko, provide minimum dates for the last deglaciation suggesting that the outlet here underwent gradual recession from 12.4 to 10.3 ka b.p. (fig. 22). the role of the inner part of egedesminde dyb, which has depths in excess of 1000 m, is unclear. the available descriptions and sea charts do not reveal a regular, u-formed conduit, but the western, steep basalt wall shows characteristics indicative of glacial plucking. we suggest that it was only the site of a major ice stream that helped to drain disko bugt during the initial thinning of the ice cover for a short time interval around 10.5–10 ka b.p. it is certain that by c. 10 ka b.p. a major break-up of the disko bugt ice cover had taken place, and this probably occurred very rapidly, over a century or less. it is presumed that the change from the large, marine-based piedmont ice lobe to a land-based ice margin, with consequent changes in geometry, resulted in the prolonged halt or slowdown of ice-margin recession in the central parts of disko bugt that lasted from c. 9.9 to c. 7.9 ka b.p. as long as the front was situated close to isfjeldsbanken near ilulissat, the calfice production of jakobshavn isbræ was probably much reduced. the cold event around 8.2 ka b.p. may also have contributed to the low recession rate. developments were different in the northern part of disko bugt, where the torsukattak icefjord, the eastern continuation of the vaigat strait, experienced a more continuous recession that brought the ice margin close to its present position before 8 ka b.p. however, the presence of moraines indicates minor deceleration or halts in the recession, perhaps at pinning points. an iceberg bank similar to that near ilulissat occurs at the junction between the torsukattak icefjord and the vaigat strait. the ice margin was presumably situated here from around 9.9 ka b.p., contemporaneous with the initial ice-margin position at the jakobshavn iceberg bank. however, the ice margin was only situated at the mouth of torsukattak icefjord for a short time, and also had a reduced frontal area. the relatively deep fjord may have favoured the continuous recession. around jakobshavn isbræ, the ice margin receded to its present location somewhat later, with 6.1 ka b.p. as a minsummary and outlook bulletin 14: gsb191-indhold 04/12/07 14:37 side 65 66 imum age; recession of the ice-sheet margin continued to the east during the holocene thermal maximum. large undocumented fluctuations of jakobshavn isbræ may have occurred during the neoglacial. we suggest that the high calf-ice production of jakobshavn isbræ began after the initial recession from isfjeldsbanken, after c. 8 ka b.p. after the subsequent recession, at 5–4 ka b.p., the frontal position was at least 15–20 km east of the present location. south of disko bugt, a relatively fast recession over the lowlands ended with the ice margin receding to its present position probably as early as 8 ka b.p. in this region, recession of the ice margin also continued beyond (east of ) the present location. the extent of marine deposits in this region points to a partially marine ice margin during most of the recession. the depths of the fjords in this region, as far as is known, do not indicate the presence of troughs that could support major ice streams during the recession. ice streams are normally located in the ablation area of the ice sheet over greenland, at sites where sufficiently deep troughs in the subsurface can act as conduits for the ice streams. with the gradual recession of the ice margin during the wisconsinan to holocene transition, periodic formation of ice streams might be expected during recession. the temporary role of such ice streams with respect to the total mass balance of the ice-sheet sector draining into the disko bugt region has still to be evaluated. detailed mapping of the entire subsurface beneath the ice margin is essential in constructing the history of development of the ice sectors of the disko bugt region. if a prerequisite for the development of an ice stream is that it is located in a deep conduit in the marginal areas of the ice sheet, the life of jakobshavn isbræ with its present activity may be restricted to the holocene period since c. 8 ka b.p. the subsequent little ice age readvance culminated in the area around jakobshavn isbræ with a major readvance in the middle and late 19th century. the extent of this readvance is unknown for the surrounding areas, with the exception of the paakitsoq area. historical information for the 20th century shows a marked recession only around jakobshavn isbræ, whereas the other ice-sheet sectors draining to disko bugt show a quasi-stability or even a tendency to advance over this period. during the holocene thermal maximum and the subsequent cooling, jakobshavn isbræ controlled much of the ice drainage of central west greenland, and its marked sensitivity to temperature changes must be linked to ablation and sliding mechanisms at the base of the ice-sheet margin. in general, disko bugt has played an important role for ice drainage of the central western slope of the ice sheet since the wisconsinan. however, the role of individual factors such as ice streams, general ablation and related subglacial meltwater transfer to the bottom of the ice and marine versus land-based ice margin, still needs to be quantified. this can be achieved through better mass-balance investigations and dynamic modelling of the ice-margin change through time. furthermore, the role of the individual troughs in disko bugt and offshore as conduits for the ice should be considered. the scattered positions of these troughs, and their relationship to temporary halts of the ice margin, point to a shift in the nature and position of ice streams during the recession of the ice margin since the last glacial maximum. whether these offshore troughs had the same central role as the present ice stream of jakobshavn isbræ during the recession is still to be docu mented. acknowledgements much of this compilation was originally made for the nomination of ‘ilulissat icefjord’ as a world heritage site. we are grateful to naja mikkelsen (geus), who co-ordinated the nomination project. we wish to thank andreas ahlstrøm, michelle citterio, robert s. fausto, almut iken, niels tvis knudsen, christoph mayer, frank nielsen, steffen pollech, niels reeh, cindy starr, henrik højmark thomsen and jacob c. yde for glaciological advice and help, james a. chalmers, t. chris r. pulvertaft, peter japsen and troels f.d. nielsen for advice and discussions on the prequaternary geology, and torben bidstrup, antoon kuijpers and naja mikkelsen for discussions on the marine geology. the manuscript has benefited from reviews by adam a. garde, niels reeh, a.k. higgins and w. stuart watt; 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giovinetto, m.b. 2001: balance, mass flux and ice vel ocity across the equilibrium line in drainage systems of greenland. journal of geophysical research 106(d24), 33,717–33,728. zwally, h.j., abdalati, w., herring, t., larson, k., saba, j. & steffen, k. 2002: surface melt-induced acceleration of greenland ice-sheet flow. science 297(5579), 218–222. bulletin 14: gsb191-indhold 04/12/07 14:37 side 75 76 aasiaat (egedesminde) town akuliarutsip sermersua (nordenskiöld gletscher) glacier akulliit island alanngorliup sermia glacier alluttoq (arveprinsen ejland) island ammassalik (tasiilaq) town, east greenland arfersiorfik fjord arveprinsen ejland (alluttoq) island buksefjorden (kangerluarsunnguaq) fjord camp century former ice-sheet station christianshåb (qasigiannguit) town claushavn (ilimanaq) town disko (qeqertarsuaq) island disko banke shallow offshore area disko bugt (qeqertarsuup tunua) bay disko fjord (kangerluk) fjord dye 3 former ice-sheet station egedesminde (aasiaat) town egedesminde dyb submarine trough eqalorutsit killiit sermiat glacier eqaluit (laksebugt) bay eqip sermia glacier gade gletscher glacier godhavn (qeqertarsuaq) town godthåb, (nuuk) capital of greenland gunnbjørn fjeld mountain, east greenland harald moltke bræ (ullip sermia) glacier hareøen (qeqertarsuatsiaq) island helheimgletscher glacier, east greenland hellefisk-1 offshore well ikerasaap sullua (qarajaq isfjord) icefjord ilimanaq (claushavn) town ilulissat (jakobshavn) town ilulissat icefjord icefjord (world heritage site) inland ice greenland icecap innaarsuit land area isfjeldsbanken submarine threshold jakobshavn (ilulissat) town jakobshavn isfjord (kangia) icefjord jakobshavn isbræ (sermeq kujalleq) glacier jøkelbugten bay kane basin large bay (basin) kangerluarsuk (vaskebugt) bay kangerluarsunnguaq (buksefjorden) fjord kangerlussuaq (søndre strømfjord) fjord/airport kangerlussuaq fjord, east greenland kangerlussuaq branch of uummannaq fjord kangerlussuaq gletscher glacier, east greenland kangersooq (nordfjord) fjord kangersuneq fjord kangia (jakobshavn isfjord) icefjord kangilerngata sermia glacier kangilliup sermia (rink isbræ) glacier kap farvel (nunap isua) cape laksebugt (eqaluit) bay ‘lerbugten’ bay lersletten (naternaq) clay plain ‘lersletten’ (narsarsuaq) clay plain marraq clay plain melville bugt (qimusseriarsuaq) bay nanortalik town narsap sermia glacier narsarsuaq (‘lersletten’) at ilimanaq clay plain narsarsuaq airport, south greenland naternaq (lersletten) clay plain nordenskiöld gletscher (akuliarutsip sermersua) glacier northgrip ice-core site nordfjord (kangersooq) at disko island fjord nuuk peninsula nuuk (godthåb) capital of greenland nuussuaq peninsula appendix 1 index to greenland place names names in quotation marks are informal names. unless stated otherwise, the localities are in west greenland. bulletin 14: gsb191-indhold 04/12/07 14:37 side 76 77 oqaatsut (rodebay) settlement orpissooq fjord paakitsoq bay paamiut (frederikshåb) town pattorfik coastal stretch peary land region, north greenland pinguarsuit rock knoll pituffik (thule air base) airbase qajaa ruin site qapiarfiit land area qaqortoq (julianehåb) town qarajaq isfjord (ikerasaap sullua) icefjord qasigiannguit (christianshåb) town ‘qarsortoq’ coastal cliff qeqertarsuaq (disko) island qeqertarsuaq (godhavn) town qeqertarsuatsiaq (hareøen) island qeqertarsuatsiaq island south of aasiaat qeqertarsuup tunua (disko bugt) bay qilertinnguit mountain rink isbræ (kangilliup sermia) glacier rodebay (oqaatsut) settlement ‘sandbugten’ bay sanddalen valley saqqarliup sermia glacier serfarsuit headland, kangersuneq fjord sermeq (upernavik isstrøm) glacier sermeq avannarleq in kangia glacier sermeq avannarleq in torsukattak glacier sermeq kujalleq (jakobshavn isbræ) glacier sermeq kujalleq in torsukattak glacier sermeq kujalleq (store gletscher) in qarajaq icefjord glacier sermermiut bay/ruin site sigguup nunaa (svartenhuk halvø) peninsula sikuiuitsoq near ilulissat fjord sisimiut (holsteinsborg) town søndre strømfjord (kangerlussuaq) fjord/airport store gletscher (sermeq kujalleq in qarajaq icefjord) glacier store hellefiskebanke shallow offshore area storstrømmen glacier sullorsuaq (vaigat) strait summit ice-sheet station svartenhuk halvø (sigguup nunaa) peninsula swiss station ice-sheet station tasiusaq fjord thule air base (pituffik) airbase tininnilik ice-dammed lake tissarissoq former part of ice-sheet margin torsukattak icefjord tuapaat coastal stretch upernavik town upernavik isstrøm (sermeq) glacier ussuit bay uummannap kangerlua (uummannaq fjord) fjord uummannaq fjord (uummannap kangerlua) fjord vaigat (sullorsuaq) strait vaskebugt (kangerluarsuk) fjord bulletin 14: gsb191-indhold 04/12/07 14:37 side 77 78 radiocarbon analyses new radiocarbon ages reported in tables 3 and 4 were determined by accelerator mass spectrometry (ams). dating was performed at the ångström laboratory at uppsala, sweden, under the supervision of göran possnert. the outer part of the shell material was removed by hydrochloric acid (hcl) to prevent contamination. the ages are reported in conventional radiocarbon years b.p. (before present = a.d. 1950). the dates have been corrected for isotopic fractionation by normalising to a δ13c value of –25‰ on the pdb scale. the δ13c measurements were performed on a conventional mass spectrometer. the dates have been corrected for a seawater reservoir effect by using an apparent age of 400 years (bennike 1997). the reservoircorrected ages have been calibrated into calendar years before present (= a.d. 1950) using the intcal04 dataset and the oxcal version 3.10 software program (bronk ramsey 2001). with respect to previous radiocarbon age determinations, compiled in tables 2–4, those marked aar and aa were also determined by accelerator mass spectrometry, whereas the other analyses were carried out by conventional methods. older dates on marine material have been reservoir corrected by subtracting 400 years (laboratory codes aar, aa) or no corrections were applied (laboratory codes i, k, hel). the latter dates have been corrected for isotopic fractionation by normalising to a δ13c value of 0‰ on the pdb scale, or no correction for isotopic fractionation was applied. these dates have a ‘built-in’ correction. appendix 2 bulletin 14: gsb191-indhold 04/12/07 14:37 side 78 79 bulletin 14: gsb191-indhold 04/12/07 14:37 side 79 80 bulletin 14: gsb191-indhold 04/12/07 14:37 side 80 research article voutchkova et al. 2021: geus bulletin 47. 6090. https://doi.org/10.34194/geusb.v47.6090 1 of 16 estimating pesticides in public drinking water at the household level in denmark denitza d. voutchkova*1 , jörg schullehner1,2 , carina skaarup3 , kirstine wodschow3 , annette kjær ersbøll3 , birgitte hansen1 1geological survey of denmark and greenland (geus), aarhus, denmark, 2department of public health, research unit for environment, work and health, aarhus university, aarhus, denmark, 3national institute of public health, university of southern denmark, copenhagen, denmark abstract pesticide pollution has raised public concern in denmark due to potential negative health impacts and frequent findings of new substances after a recent expansion of the groundwater monitoring programme. danish drinking water comes entirely from groundwater. both the raw groundwater and the treated drinking water are regularly monitored, and the chemical analyses are reported to a publicly available national database (jupiter). based on these data, in this study we (1) provide a status of pesticide content in drinking water supplied by public waterworks in denmark and (2) assess the proportion of danish households exposed to pesticides from drinking water. ‘pesticides’ here refers also to their metabolites, degradation and reaction products. the cleaned dataset represents 3004 public waterworks distributed throughout the country and includes 39 798 samples of treated drinking water analysed for 449 pesticides (971 723 analyses total) for the period 2002–2019. of all these chemical analyses, 0.5% (n = 4925) contained a quantified pesticide (>0.03 µg/l). pesticides were found at least once in the treated drinking water at 29% of all sampled public waterworks for the period 2002–2019 and at 21% of the waterworks for the recent period 2015–2019. we estimate that 56% of all danish households were potentially exposed at least once to pesticides in drinking water at concentrations of 0.03–4.00 µg/l between 2002 and 2019. however, in 2015–2019, the proportion of the danish households exposed to pesticides (0.03–4.00 µg/l) was 41%. the proportion of danish households potentially exposed at least once to pesticides above the maximum allowed concentration (0.1 µg/l) according to the eu drinking water directive (and the danish drinking water standard) was 19% for 2002–2019 and 11% for 2015–2019. however, the maximum concentrations were lower than the world health organization’s compound-specific guidelines. lastly, we explore data complexity and discuss the limitations imposed by data heterogeneity to facilitate future epidemiological studies. 1 introduction pesticides are biologically active compounds widely used in agriculture, horticulture and public health for the control of pests (world health organization 2019a). they comprise many chemical substances with a broad variety of mode of action depending on their target organisms (casida 2009), for example, photosynthesis inhibition (plants), neurotoxic (insects) and fungal spore *correspondence: dv@geus.dk received: 24 nov 2020 accepted: 10 feb 2021 published: 12 apr 2021 keywords: denmark, drinking water, exposure, pesticides, public waterworks abbreviations bam: 2,6-dichlorobenzamide dbcp: dibromochloropropane deia: desethyl-desisopropyl atrazine dpc: desphenyl chloridazon dms: n,n-dimethylsulfamide ddt: dichlorodiphenyltrichloroethane dwd: drinking water directive dwqs: drinking water quality standard epa: environmental protection agency geus: geological survey of denmark and greenland mdpc: methyl-desphenyl-chloridazon lod: limit of detection loq: limit of quantification who: world health organization wsa: water supply area geus bulletin is an open access, peerreviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: adam hambly (technical university of denmark) reviewed by: martin rygaard (technical university of denmark), kai tang (technical university of denmark) funding: see page 14 author contributions: see page 14 competing interests: none declared additional files: see page 15 https://doi.org/10.34194/geusb.v47.6090� https://orcid.org/0000-0003-2840-072x https://orcid.org/0000-0002-1153-6885 https://orcid.org/0000-0001-7470-5052 https://orcid.org/0000-0001-9908-3632 https://orcid.org/0000-0002-9407-3387 https://orcid.org/0000-0003-2318-145x mailto:dv@geus.dk voutchkova et al. 2021: geus bulletin 47. 6090. https://doi.org/10.34194/geusb.v47.6090 2 of 16 www.geusbul let in.org inhibition (fungi). for a classification of over 180 active substances (pesticides) and their metabolites based on the mode of action, see mohaupt et al. (2020; annex 5). agricultural pesticides threaten aquatic biodiversity (stehle & schulz 2015) and potentially impact food-webs and species competition (kohler & triebskorn 2013). new pesticides are registered for use only if they are demonstrated not to persist in the environment after their intended period of use. however, residues are found in the natural environment in nanogram to microgram per litre levels (fenner et al. 2013). for example, the status assessments for water bodies in the european environmental agency countries showed that 0.4% of all surface water bodies and 6.5% of the area of groundwater bodies are failing ‘good chemical status’ because of pesticides (mohaupt et al. 2020). in fact, pesticides are second only to nitrates as a cause for failing to achieve a level of good chemical status in europe (mohaupt et al. 2020). this shows that pesticides persist in the environment, and the only natural removal is through degradation by biotic or abiotic transformation processes (fenner et al. 2013). 1.1 pesticide exposure and drinking water pesticides are inherently hazardous compounds, but even though only a small fraction is highly hazardous (world health organization 2019b), they cause disproportionate harm to the environment and human health worldwide (food and agriculture organization & world health organization 2019). the greatest human exposure to pesticides occurs in occupational environments, for example, in factory, agricultural and public health workers during production, handling, dilution, mixing and application procedures (world health organization 2019a). acute pesticide poisoning is an important cause of mortality and morbidity, for example, due to neurotoxic effects of organophosphate intoxication (rosenstock 1991). however, there is incomplete knowledge on the toxicity of the metabolites, degradation and reaction products, which could have a similar, stronger or lesser effect on organisms and humans (mohaupt et al. 2020; p. 7, box 1). the general population may be environmentally exposed to pesticide residues from food and drinking water (world health organization 2019a). a variety of chronic health effects related to exposure at doses that do not cause acute effects have been suggested, including asthma, diabetes, parkinson’s disease and cancer (kim et al. 2017). protecting the population from health risks associated with pesticide-contaminated drinking water is a worldwide problem (li & jennings 2018). thirty-four percent of the world’s population, in >1/2 of the world’s nations, are estimated to be inadequately protected against health risks associated with pesticide-contaminated drinking water (li & jennings 2018). pesticide pollution was recently found to be of major concern in drinking water sources in the netherlands, where pesticides were found in 2/3 of the water abstraction areas. the water-quality standard was exceeded in 1/3 of all drinking-water sources in the netherlands (including both groundwater and surface-water bodies; sjerps et al. 2019). bexfield et al. (2021) also showed that at least one pesticide or degradate was found in 41% of wells investigated in the united states (n = 12 041 204 in aquifers responsible for 70% of the volume pumped for public drinking-water supply, nationally), and around 2/3 of them contained compound mixtures. although pesticide compounds occurred frequently, concentrations were low, and only 1.6% of wells had concentrations approaching levels of potential health concern (bexfield et al. 2021). pesticide exposure from drinking water in denmark is estimated to be smaller than that from other dietary sources, such as berries, fruits and vegetables (bichel hovedudvalget 1999). nevertheless, concerns about adverse health outcomes due to long-term low-dose intake from drinking water have been raised, and there has not been enough scientific evidence to either support or reject potential health risks (bichel hovedudvalget 1999). while the pesticides found in food products are mainly contemporary insecticides and fungicides (approved for use), the pesticides found in drinking water are mainly herbicides, often representing legacy pollution. the term ‘pesticide’ covers a large group of substances with different structure and mechanism of action, making it relevant to study the presence and potential influence of pesticides in drinking water. to the best of our knowledge, there are no recent publications on drinking water as a dietary source of pesticides in denmark. exposure estimates from drinking water in denmark are thus lacking. 1.2 objectives the purpose of this article is (1) to assess the pesticide status of danish drinking water, (2) to estimate the population’s exposure to pesticides from drinking water and (3) to evaluate the dataset heterogeneity and limitations. we take advantage of the national database (jupiter; https://eng.geus.dk/products-services-facilities/data-andmaps/national-well-database-jupiter) in which all drinking water samples taken for compliance or other purposes are centrally registered. the study period is 18 years (2002–2019) with sufficient data coverage to analyse the current drinking water status at a national level for the entire study period and specifically, the last five years (2015–2019). https://doi.org/10.34194/geusb.v47.6090� www.geusbulletin.org� https://eng.geus.dk/products-services-facilities/data-and-maps/national-well-database-jupiter� https://eng.geus.dk/products-services-facilities/data-and-maps/national-well-database-jupiter� voutchkova et al. 2021: geus bulletin 47. 6090. https://doi.org/10.34194/geusb.v47.6090 3 of 16 www.geusbul let in.org 2 study setting and legal framework european union member states must ensure that water intended for human consumption is ‘wholesome and clean’, according to the drinking water directive (dwd; council of the european union 2015). this means that drinking water should comply with the dwd’s minimum requirements for several microbiological and chemical components, including pesticides. the maximum allowed concentration of pesticides in drinking water – the drinking water quality standard (dwqs) – is set in the dwd at (1) 0.10 µg/l for individual compounds, except for aldrin, dieldrin, heptachlor and heptachlor epoxide, which are set at 0.030 μg/l, and (2) 0.50 μg/l for the sum of detected and quantified pesticides (named ‘pesticides – total’ in the dwd). the dwd also provides a legal definition of the term ‘pesticide’, which we follow in this article: ‘pesticides means [-] organic insecticides, [-] organic herbicides, [-] organic fungicides, [-] organic nematocides, [-] organic acaricides, [-] organic algicides, [-] organic rodenticides, [-] organic slimicides, [-] related products (inter alia, growth regulators) and their relevant metabolites, degradation and reaction products’. (annex i, part b, note 6 of the dwd; the square brackets indicate a change in the punctuation from the original dwd) denmark has implemented the dwd provisions in national legislation (vandforsyningsloven; lbk nr 118, 22 february 2018). actionable items (e.g. sampling frequencies and dwqs) are further specified by ministerial orders (the most recent is bek nr 1070, 28 october 2019). the frequency of sampling varies based on the volumes of drinking water produced, from once every 3 years to several times per year (bek nr 1070, 28 october 2019). the list of compounds monitored for mandatory compliance is revised every year, and currently, it contains 46 pesticides. waterworks must also test for other pesticides if there is evidence that their presence in the catchment area poses a health threat. danish drinking water supply is decentralised (>4500 public waterworks) and 100% groundwater based. it mostly relies on simple groundwater treatment, including aeration and sand filtration. if the simple treatment is not sufficient to assure compliance with the dwqs, the municipalities can grant permits for the use of advanced treatment processes. in the period 2007–2012, 74 waterworks obtained such permits, but of those, only eight (representing c. 2.5% of treated groundwater) were for carbon filters to treat for 2,6-dichlorobenzamide (bam) or other organic micropollutants (naturstyrelsen 2012). in the period 2012–2019, the number of permits granted for advanced treatment increased to 110, of which 12 were for carbon filters (miljøstyrelsen 2020a). in addition to advanced treatments, waterworks could address non-compliance by closing polluted wells or well fields, diluting non-compliant water with groundwater from another well or importing unpolluted water from neighbouring waterworks. all laboratory results of both drinking water and groundwater samples in denmark are reported to the danish national well database (jupiter), according to the national guidelines (miljøstyrelsen 2020b). based on these data, the geological survey of denmark and greenland (geus) has reported the status of raw groundwater in denmark on an annual basis for the past 30 years (thorling et al. 2019). additionally, the danish environmental protection agency (epa) conducted a mass screening in 2019, where 263 groundwater wells were tested for 415 pesticides (mathiesen 2020). based on these tests, the danish epa planned to revise the list of pesticides for mandatory compliance monitoring of drinking water and to intensify its control efforts against illegal imports of pesticides (mathiesen 2020). our study provides a national assessment of pesticides in danish treated drinking water, supplied by public waterworks and on that basis an estimation of households exposed to pesticides from drinking water. 3 methods and materials the complete workflow used in this study is presented schematically in fig. 1. a similar exposure-estimation methodology was used previously (schullehner & hansen 2014; voutchkova et al. 2015). here, we outline all data-handling procedures to assure reproducibility and transparency, and to facilitate the potential future use of the drinking water dataset. 3.1 data sources the final dataset (fig. 1) used for the status overview and to estimate pesticide exposure was prepared by combining data from four different sources. 3.1.1 drinking water samples analysed for pesticides all chemical analyses of drinking water reported to jupiter were extracted on 5 may 2020 (fig. 1, step 1). this raw dataset was passed through a pre-processing procedure including various filtering and quality control steps summarised here and presented in full detail in supplementary file s1. compound selection was based on the latest version (20 may 2020) of the jupiter list of pesticides, their degradation products and related substances (‘50 pesticider, nedbrydningsprodukter og beslægtede stoffer’ in jupiter). only samples of treated drinking water taken at the waterworks (i.e. the finished product), from the distribution network or the consumer’s tap were kept in https://doi.org/10.34194/geusb.v47.6090� www.geusbulletin.org� voutchkova et al. 2021: geus bulletin 47. 6090. https://doi.org/10.34194/geusb.v47.6090 4 of 16 www.geusbul let in.org the dataset (see supplementary file s1, section 1.1, step 4 for details). we filtered out flagged or erroneous data and those data rejected by the data-owner. analytical units were checked and converted when necessary to µg/l. all analyses below detection limit in instances where that detection limit was high (>0.02 µg/l) were excluded due to low precision of the laboratory method. after the pre-processing procedure fig. 1 workflow including data-processing steps and overview of excluded chemical analyses and waterworks. dk: denmark. wsa: water supply areas. qc: quality control and filtering procedures (see text for details). plant_ id and wsa_id: waterworks and wsa id numbers, respectively, used when combining the datasets. sf: supplementary files. jupiter 1 003 804 analyses 3173 waterworks 449 compounds x & y coordinates public dk waterworks (n = 4641) 989 218 analyses 3 072 waterworks 449 compounds plant_id 2584 water supply areas for 4456 waterworks qc1 8 151 264 analyses 4411 waterworks 1140 compounds 5 may 2020 join qc2 14 586 analyses 101 waterworks details in sf join17 495 analyses 68 waterworks final pesticides dataset no wsa no x,y status 1 individual compounds status 2 waterworks status 3 water supply areas 99.46% of dk households for 2811 wsa exposure to pesticides join wsa_id status and exposure calculations were done: 1) over the entire study period (2002–2019) 2) for the last five years (2015–2019) no households in 4 wsa step 1 step 2 step 3 plant_id water supply areas geocoded households 971 723 analyses 3004 waterworks 449 compounds database processing step process output other input datasets symbols https://doi.org/10.34194/geusb.v47.6090� www.geusbulletin.org� voutchkova et al. 2021: geus bulletin 47. 6090. https://doi.org/10.34194/geusb.v47.6090 5 of 16 www.geusbul let in.org (supplementary file s1, section 1.1), the resulting dataset contained all pesticide analyses of treated drinking water for the period 2002–2019 from public waterworks (n = 3173; fig. 1, step 2). 3.1.2 geographic coordinates of the waterworks the geographic coordinates of all waterworks were extracted from jupiter and passed through a quality control. missing x and y coordinates were geocoded based on the registered address. coordinate errors were identified and fixed when possible. only coordinates for public waterworks (n = 4641) were kept, from which 5.5% (n = 256) were geocoded based on address. 3.1.3 water supply areas of the waterworks the water supply areas (wsas) for all public waterworks, covering the entire country, were provided by schullehner and hansen (2014). we assumed that wsas were static in the study period (2002–2019), reflecting their state at the time the data were published (2014). 3.1.4 percent of households located within a wsa geocoded locations of households were provided by the centre for integrated register-based research at aarhus university (cirrau), denmark. this data source contains all residential addresses in denmark registered in the danish civil registration system (n = 2 086 797; pedersen 2018). just 0.54% of all households fell outside a polygon for the wsa. 3.2 data joining procedure data from the four sources were combined using a stepwise procedure (‘join’ in fig. 1). the pre-processed pesticide dataset from step 2 was joined with the waterworks’ coordinates. in this step, we excluded 101 waterworks with all their analyses. most of these were waterworks located in greenland (n = 70), and four were identified as misclassified private waterworks. the remaining waterworks (n = 27) had no coordinates or a registered address and could not be geocoded. of these, only four were actually waterworks, and the rest were locations at the supply network or at the consumers’ tap. then, the dataset (fig. 1, step 3) was joined with the wsa data, resulting in the exclusion of 68 waterworks which could not be assigned a wsa. we refer to this dataset as the final pesticide dataset. 3.3 final pesticide dataset the final pesticide dataset includes 39  798 treated drinking-water samples (7 january 2002–30 dec 2019) analysed for 449 pesticides (number of individual analyses n = 971  723), associated with public waterworks located in denmark (n = 3004) with known x and y coordinates and wsa. figure 2 shows the spatial distribution of these waterworks, the number of waterworks within the wsa and the percent of households in each wsa. the final dataset was used further to produce the status overviews of individual pesticides, waterworks and wsas (status 1, 2 and 3 in fig. 1). the population exposure to pesticides from drinking water was obtained by combining status 3 results (fig. 1) with the percentage of households with each wsa (fig. 2). the methodology for these status overviews and the exposure assessments are presented in section 3.4. first, we present the methodology for how we handled values below the detection/quantification limit and discuss the data structure. 3.3.1 limit of detection and quantification the laboratories performing the chemical analyses report the limit of detection (lod) in the jupiter database. all measurements below the lod are recorded with the attribute ‘<’ and a value equal to the lod (e.g. <0.01 µg/l). most of the pesticide analyses in the final dataset were below the specified lod (98.8%, n = 960 437). lod varied from 0.4 ng/l to 0.02 µg/l. analyses with higher lods were excluded in the pre-processing. the most frequent lod was 0.01 µg/l (99.2%, n = 953 120), followed by 0.02 µg/l (0.6%, n = 5522; see supplementary file s1). the variation in lods reflects the variety of methods used by different laboratories, with liquid chromatography–mass spectrometry (lc/ms/ms) being the most frequently used method (see supplementary file s1). the variation with time is also due to possible instrumentation improvements, reducing the lod during the study period. in the status assessments and the exposure estimation, we use the limit of quantification (loq) instead of the lod to avoid the effect of false positive detections and to reduce the influence of the high uncertainty in concentrations measured near the lod. evidence of the variability around the lod is presented in section 3.3.2. the loq was calculated according to the ministerial order on quality requirements for environmental measurements (bek nr 1071, 28 october 2019) with the formula loq = 3 × lod, where lod here is the most frequent lod (0.01 µg/l). thus, loq = 0.03 µg/l (equal to 30% of the dwqs). for all analyses with concentrations less than loq, we adopt the definition that ‘pesticides were not determined with an acceptable level of accuracy and precision’ (commission of the european communities 2009). for brevity, we use ‘non-detect’ or ‘never detected’, where this definition applies to a period of time or pesticide, respectively. https://doi.org/10.34194/geusb.v47.6090� www.geusbulletin.org� voutchkova et al. 2021: geus bulletin 47. 6090. https://doi.org/10.34194/geusb.v47.6090 6 of 16 www.geusbul let in.org 3.3.2 structural heterogeneity the final dataset structure is characterised by many types of heterogeneity. here we illustrate those caused by (1) the dynamic nature of the analytical program, (2) the varying sampling frequency and consequently the varying length of the sampling gaps and (3) the variability of the analytical results around the lod. inspecting the timeseries of individual pesticides at various waterworks shows that it is not uncommon to have concentrations varying around the lod. this includes frequent changes between detects (concentrations ≥lod) and non-detects (concentrations <lod), sometimes within the same sampling event (see supplementary file s1 for examples). we quantified this instability by counting the number of changes from <lod to ≥lod for each waterworks and each pesticide. figure 3 shows the proportion of waterworks where such changes occurred for individual pesticides. the least stable results around the lod are for bam – about 15% (n = 469) of all waterworks with bam analyses in the study period (n = 2997) had multiple changes. bentazon (n = 89), dichlorprop (n = 63), atrazine-desisopropyl (n = 56) and desphenyl chloridazon (dpc; n = 53) are in the top five based on the absolute number of waterworks with multiple changes around the lod. the dynamic nature of the analytical programme for pesticides is reflected in the data availability for each pesticide in the final dataset. the period covered for individual pesticides varies from 1 to 18 years (2002– 2019); 64.8% of all 449 pesticides were only analysed in 1 year, while 14.7% were analysed throughout the entire study period (fig. 4a). many pesticides (71%) were analysed for the first time in 2018 or 2019 (fig. 4b). the latest sampling date for almost all pesticides was in 2019 (n = 432, 96.2%). the third heterogeneity aspect relates to the sampling frequency. the number of sampling events at the fig. 2 number of public waterworks with pesticide data for each water supply area (wsa) for a: the entire study period (2002–2019) and b: the last five years (2015–2019). c: percentage of danish households (n = 2 086 797) located in each of the wsa, equal count in each interval. d: location of all public waterworks with at least one pesticide analysis in the period 2002–2019 included in the final dataset. note: the symbols in d have transparency to visualise overlapping locations. na: no data. 1 2 3 4 5 6 7 8 9 11 na number of waterworks with pesticides data per wsa in the period 2002–2019 (n) n n households in a wsa (%) 0.000–0.002 0.002–0.005 0.005–0.009 0.009–0.017 0.017–0.044 0.044–1.423 n a 1 2 3 4 5 6 8 na number of waterworks with pesticides data per wsa in the period 2015–2019 (n) locations of waterworks with pesticide data 2002–2019 b n c d https://doi.org/10.34194/geusb.v47.6090� www.geusbulletin.org� voutchkova et al. 2021: geus bulletin 47. 6090. https://doi.org/10.34194/geusb.v47.6090 7 of 16 www.geusbul let in.org 3004 waterworks varies from once to >200 times within the study period (2002–2019) with the most common being nine times (7.9% of the waterworks; fig. 5a). of all waterworks, 4.6% had only one sampling event in that period. in the past five years, there were 2405 waterworks with data, of which 7.6% had only one sampling event. the most common sampling frequency was 3 times (at 19.9% of the waterworks; fig. 5b). because of the variation in sampling frequency, the data gaps also vary. figure 5c and d shows the mean sampling gap (the number of days between sampling events) for 2002– 2019 and 2015–2019, respectively. 3.4 status assessment and exposure estimation the status assessment is presented in four qualitative groups based on the determined pesticide concentrations as follows: (1) never detected (<loq, where loq = 0.03 µg/l), (2) detected but always ≤0.1 µg/l, (3) exceeded 0.1 µg/l (the dwqs for individual pesticides) and (4) not analysed (no data). a pesticide had to be detected or exceed the dwqs at least once within the period of interest (status 2) and for at least one waterworks in the wsa (status 3). where a wsa contained multiple waterworks (fig. 2), the entire area was classified as exceeding the dwqs if at least one exceeded the dwqs. the exposure to pesticides from drinking water was then estimated for the following four exposure classes: (1) never exposed to pesticides from drinking water, (2) potentially exposed to low concentration of pesticides (0.03–0.1 µg/l), (3) potentially exposed to concentrations exceeding the dwqs (>0.1 µg/l) and (4) unknown exposure due to no data. these general classes of exposure were selected because of the structural heterogeneity of the final dataset (section 3.3.2). our methodology is in line with the deterministic approaches for estimating exposure, which usually include worst-case assumptions and result in a conservative estimate of exposure (ferrier et al. 2002). fig. 3 proportion of waterworks with ‘unstable’ detections of pesticides during the study period (2002–2019). ‘unstable’ is defined here as having more than one fluctuation around the lod (i.e. multiple changes from <lod to ≥lod or vice versa). data labels in the format n / n indicate number of waterworks with unstable detections versus all waterworks with pesticide data. pesticides labels are used as provided in jupiter database. https://doi.org/10.34194/geusb.v47.6090� www.geusbulletin.org� voutchkova et al. 2021: geus bulletin 47. 6090. https://doi.org/10.34194/geusb.v47.6090 8 of 16 www.geusbul let in.org status 1 for individual pesticides (fig. 1) encompasses the entire study period (2002–2019), while the status of waterworks and wsas (status 2 and 3) and the exposure estimates were prepared for both 2002–2019 and 2015– 2019. by using two periods, we can account for potential changes due to the measures taken by waterworks to comply with the dwqs limits. only waterworks with a minimum of one pesticide analysis in the respective period are included in these calculations. the methodological details for status assessments for the sum of fig. 4 data availability. a: period (in years) where data are available for individual pesticides, calculated as the difference between first and last sampling date for each of the 449 compounds. b: the year when individual pesticides were first measured. data labels show the number of individual pesticides. fig. 5 frequency of sampling events by sampling date in a: 2002–2019 and b: 2015–2019. the cumulative proportion of the mean sampling gap per waterwork (period between two sampling events) in c: 2002–2019 and d: 2015–2019. c and d include only waterworks with more than one sampling event (n = 2866 for c; n = 2222 for d). https://doi.org/10.34194/geusb.v47.6090� www.geusbulletin.org� voutchkova et al. 2021: geus bulletin 47. 6090. https://doi.org/10.34194/geusb.v47.6090 9 of 16 www.geusbul let in.org pesticides (i.e. ‘pesticides – total’ in the dwd) are provided in supplementary file s1, section 2.4. 3.5 software all data processing and analyses were done in r v.3.6.0 (r core team 2019) and rstudio desktop 1.2.1335 (rstudio team 2018) with additional r packages (wickham 2016; dowle & srinivasan 2019; r core team 2019; wickham & henry 2019; wickham et al. 2019). qgis desktop v.3.6.1 (qgis development team 2019) and inkscape v.0.92.4 (inkscape project 2019) were used to produce the maps and flowchart, respectively. 4 results 4.1 status overview 1 – individual compounds a total of 99.5% of the individual analyses in the final dataset were <loq (n = 966 798) and 0.5% (n = 4925) contained a detected and quantified pesticide (≥0.03 µg/l). of these, 16% (n = 793) exceeded dwqs. in the period 2002–2019, 10.2% of the 449 pesticides (n = 46) were detected in at least one sample. of those, bam was analysed at most of the waterworks (n = 2997). bam is a degradation product of the banned herbicides dichlobenil and chlorthiamid, used for weed control in urban areas and in berry production in the period 1969–1996 (thorling et al. 2019). bam has been analysed throughout the entire study period and has the largest total number of analyses (n = 36 777), of which 6.8% contained a detected pesticide (n = 2514), but only 0.9% (n = 330) exceeded the dwqs (0.1 µg/l). the maximum detected concentration of bam in the study period was 1.9 µg/l. the absolute maximum detected concentration was 4 µg/l for dpc. dpc is a degradation product from the banned herbicide chloridazon, sold in denmark between 1964 and 1996 (thorling et al. 2019). dpc was analysed at 2033 public waterworks in the period 2016–2019 (except for a single analysis in 2010). from all reported dpc analyses (n = 5140), 14.4% contained a detected pesticide (n = 739), but only 4.4% (n = 227) exceeded the dwqs. next to bam and dpc, dms (n,n-dimethylsulfamide) is also in the top three of pesticides that exceeded dwqs (in terms of absolute number of analyses). dms is a degradation product from the fungicides tolylfluanid and dichlofluanid, which have been applied as both pesticides and biocides (thorling et al. 2019). dms has been analysed at 1747 waterworks since march 2018. about 18% of the 3333 dms analyses contain a detected pesticide (n = 620), and 83 (2.5%) exceeded the dwqs. aldrin, dieldrin, heptachlor and heptachlor epoxide (dwqs = 0.03 µg/l) were reported since 2018 at 1631 waterworks (aldrin, dieldrin and heptachlor) and 1600 waterworks (heptachlor epoxide). the complete status 1 overview including all 449 individual substances is provided in supplementary file s2. 4.2 status at waterworks and wsas (status 2 and 3) and pesticide exposure pesticides were detected at least once in drinking  water  (≥0.03 µg/l) at 29% of the waterworks in the  study  period (2002–2019) and at 24% of the waterworks with measurements in the last five years (2015–2019; fig.  6). about 9% of all waterworks exceeded the dwqs (>0.1 µg/l) in the study period and 7% in the past five  years. the rest of the waterworks (71–76%) had never detected and quantified a pesticide (loq = 0.03  µg/l) based on the final dataset. figure 7a and b provide a spatial visualisation of waterworks status. the status of wsas (status 3) resembles that for waterworks with minor differences in the percentages (table sm-4 in supplementary file s1). twelve percent of the wsas had at least one waterworks that exceeded the dwqs in 2002–2019, dropping to less than 8% when only the past five years were considered. the difference between the two periods can also be observed in fig. 7c and d. status 3 was used to estimate the population exposure, based on the percent of households located in each wsa. fig. 6 pesticide status at waterworks and estimated exposure at the household level. note: the classes ‘never detected’ and ‘never exposed’ refer to pesticide concentrations below the limit of quantification (loq = 0.03 µg/l). for absolute numbers, see table sm-4 in supplementary file s1. 76 71 17 20 7 9 54 42 30 37 11 19 5 2 waterworks 2002–2019 2015–2019 never detected detected ≤0.1 µg/l exceeding 0.1 µg/l 0% 20% 40% 60% 80% 100% never exposed exposed ≤0.1 µg/l exposed >0.1 µg/l unknown households 2002–2019 2015–2019 0% 20% 40% 60% 80% 100% https://doi.org/10.34194/geusb.v47.6090� www.geusbulletin.org� voutchkova et al. 2021: geus bulletin 47. 6090. https://doi.org/10.34194/geusb.v47.6090 10 of 16 www.geusbul let in.org based on the final dataset, about half of all danish households (54%) were never exposed to the pesticides analysed between 2015 and 2019 (fig. 6). this was lower (42%) for the period 2002–2019. however, fewer households were exposed to low pesticide concentrations (0.03–0.1 µg/l) in the past five years (30%) compared to the entire study period. similarly, the households potentially exposed to pesticides exceeding 0.1 µg/l were 19% for 2002–2019 and 11% for 2015–2019. for status with respect to ‘pesticides-total’ (0.5 µg/l), see supplementary file s1. we could not estimate the exposure for 2% (2002– 2019) and 5% (2015–2019) of danish households due to lack of data – 599 waterworks had no data after 2015, hence the higher percent for 2015–2019. some of these waterworks may have closed, but it is also possible that there is a delay in reporting to the danish database or that the analyses were excluded in the quality assurance procedures or the other data-preparation steps. 5 discussion 5.1 possible trends during the period 2002–2019, danish drinking water has been tested for 449 individual substances defined here as pesticides, resulting in more than 970 000 individual analyses. of which, less than 5000 (0.5%) contained a quantified pesticide. bam has the most spatially and temporally complete data coverage. a tenth of all danish households connected to public water supplies were potentially exposed to pesticides from drinking water at concentrations >0.1 µg/l and about 30% to concentrations in the range 0.03–0.10 µg/l during the past five years (2015–2019). pesticides exposure from drinking water (and the waterworks status) is lower (better) in the period 2015–2019 when compared to the entire 18-year study period. the reduction in those exposed in the past five years could be due to improved groundwater quality. also, some of the waterworks where concentrations exceeded the dwqs may have (1) closed before 2015, fig. 7. pesticide status at waterworks and water supply area (wsa). ‘never detected’ refers to all analyses or waterworks where pesticides are measured below the quantification limit (qol = 0.03 µg/l). na: no data. status at waterworks 2002–2019 never detected detected ≤0.1 µg/l * exceeded 0.1 µg/l * n status at waterworks 2015–2019 never detected detected ≤0.1 µg/l * exceeded 0.1 µg/l * n status at wsa 2002–2019 always <loq detected ≤0.1 µg/l * exceeded 0.1 µg/l * na regions n a * for at least one waterworks status at wsa 2015–2019 never detected detected ≤0.1 µg/l * exceeded 0.1 µg/l * na regions b n c * for at least one waterworks d * for at least one analysis * for at least one analysis https://doi.org/10.34194/geusb.v47.6090� www.geusbulletin.org� voutchkova et al. 2021: geus bulletin 47. 6090. https://doi.org/10.34194/geusb.v47.6090 11 of 16 www.geusbul let in.org (2) closed or changed wells and well-fields, changed their pumping strategies or started importing water from other waterworks or (3) implemented advanced treatment processes for pesticide removal. thus, this might indicate that water supply companies in denmark are working towards resolving compliance issues. however, we cannot directly compare two different periods (before/after) or do trend analysis due to the issue surrounding data heterogeneity (fig. 3). notably, dpc and dms, the two compounds most discussed in denmark today, were only analysed during the past five years. 5.2 comparison with untreated groundwater the pesticide content of untreated groundwater and treated drinking water was compared since all drinking water in denmark originates from groundwater (table 1). the status of the danish groundwater is assessed and reported annually as part of the national groundwater-monitoring program (e.g. thorling et al. 2019). however, such direct comparison between the pesticide status of treated drinking water and the raw, untreated groundwater at a national scale could be misleading. often, raw groundwater is abstracted from multiple wells and well fields and mixed before or after the treatment. it is also possible that some pesticides are removed during treatment. in the most general terms, our results for the 2015–2019 period fit well with the findings for untreated groundwater. the percent of monitoring points with at least one detected pesticide is lower in drinking water (24% for 2015–2019) than in groundwater (29 % for 2014–2018). this could be partly due to the different analytical programme and because we used a more conservative data-quality assurance. for example, we look at the detected and quantified pesticides based on loq (0.03 µg/l), while thorling et al. (2019) used the lod (0.01 µg/l). 5.3 drinking water quality standard and health implications the dwqs for pesticides in the eu is based on the ‘precautionary principle’ (bourguignon 2016). the dwqs for individual pesticides (0.1 µg/l) was a surrogate for zero when it was first introduced in 1980 (dolan et al. 2013). it remained unchanged in the 1998 revision and was regarded by dolan et al. (2013) to be still consistent with the precautionary principle. no changes to the dwqs for pesticides are planned in the revised directive (european commission 2018). the dwqs in the dwd was found to be one of the most conservative worldwide for 16 out of 25 of the most commonly regulated pesticides (li & jennings 2018). lower standards existed elsewhere for the following pesticides: lindane, chlordane, heptachlor, endrin, dibromochloropropane (dbcp), heptachlor epoxide, dichlorodiphenyltrichloroethane (ddt), aldrin and dieldrin (li & jennings 2018). the world health organization (who) drinking water guidelines, on the other hand, aim to protect against health effects resulting from a lifetime exposure (world health organization 2011, 2017). who guidelines (table 2) are established only for pesticides of significant international concern when there is credible evidence for occurrence in drinking water and actual or potential toxicity (world health organization 2017). we detected and quantified about one-third of the pesticides for which who guideline values are established, although the maximum concentrations were lower than the who guidelines (table 2). the who notes that as a general principle, concentrations of pesticides in drinking water should be kept as low as possible and increasing trends up to the guideline values should be prevented (world health organization 2017). 5.4 limitations 5.4.1 drinking water pesticide data the jupiter database is a comprehensive resource for both drinking and groundwater quality data since all laboratory samples taken from wells and waterworks table 1 results from the pesticide status of untreated groundwater (thorling et al. 2019) sampling points water type sampling period total monitoring points (n) detected1 (%) exceeded (%) monitoring wells2 groundwater 2016–2018 1084 42 18 waterworks wells3 groundwater 2014–20184 6342 29 7 waterworks5 drinking water 2015–2019 2405 24 7 the drinking water results from this study are presented for reference. 1for groundwater samples, ‘detected’ refers to concentrations ≥0.01 µg/l. in drinking water, it refers to ‘detected and quantified’ (i.e. ≥0.03 µg/l). 2all sampled intakes belonging to the monitoring wells from the danish nation-wide monitoring programme (not for drinking water production; thorling et al. 2019). 3all sampled intakes belonging to waterworks wells used for drinking water production from the analysis package ‘boringskontrol’ (thorling et al. 2019). 4all waterworks wells in denmark should have been tested for pesticides at least once in this period. the analytical programme expanded in 2018, so these percentages will most likely change once all wells are tested for the complete list of pesticides. 5from this study. https://doi.org/10.34194/geusb.v47.6090� www.geusbulletin.org� voutchkova et al. 2021: geus bulletin 47. 6090. https://doi.org/10.34194/geusb.v47.6090 12 of 16 www.geusbul let in.org table 2 drinking water guideline values for pesticides (world health organization 2017) and occurrence in danish drinking water based on our final pesticide dataset pesticide who-gv (µg/l) max (µg/l) ≥loq (n) ww (n) total analyses (n) since (year) found aldrin and dieldrin 0.03   1631 2698 2018 no chlordane 0.2   11 11 2019 no 1,2-dibromoethane 0.4      cyanazine 0.6   2906 20 015 2002 no endrin 0.6   11 11 2019 no 1,2-dibromo-3-chloropropane 1      mcpa 2 1.23 12 2900 35 153 2002 yes simazine1 2 0.051 16 2990 35 088 2002 yes lindane 2   23 35 2019 no molinate 6      dimethoate 6   2905 20 102 2002 no terbuthylazine 7 0.031 2 2801 19 657 2002 yes carbofuran 7   33 387 2002 no isoproturon 9 0.12 4 2905 20 402 2002 yes fenoprop 9      2,4,5-t 9   131 623 2002 no mecoprop 10 0.74 52 2991 35 185 2002 yes aldicarb 10   11 11 2019 no metolachlor 10 -   38 111 2018 no pendimethalin 20 0.057 3 2906 20 205 2002 yes 1,3-dichloropropene 20      alachlor2 20   38 126 2018 no methoxychlor 20   11 11 2019 no trifluralin 20   129 844 2002 no 2,4-d 30 0.08 2 2957 29 020 2002 yes chlorotoluron 30   11 11 2019 no chlorpyrifos 30   22 23 2019 no 1,2-dichloropropane 40      2,4-db 90   11 11 2019 no atrazine3 100 0.26 43 2990 35 132 2002 yes dichlorprop 100 0.87 82 2990 35 120 2002 yes hydroxyatrazine 200 0.12 5 2988 35 050 2002 yes who-gv: world health organization (who) guideline value for drinking water. max: maximum detected and quantified concentration. ≥loq: analyses above the limit of quantification (0.03 ug/l). ww: number of waterworks where the specific pesticide was analysed at least once in the study period. since: the first year of analyses (in the dataset). found: yes or no, indicating if the pesticide was found in danish drinking water. 1a simazine metabolite had higher maximum concentration (detected and quantified in 29 analyses max = 0.12 µg/l). 2alachlor’s metabolite was detected and quantified 3 times (range 0.05–0.08 µg/l). 3the guideline value also covers the chloro-s-triazine metabolites. desethyl atrazine, desisopropyl atrazine, hydroxyatrazine and desethyl-desisopropyl atrazine (deia) were also detected and quantified here. in denmark must be entered into it. however, reporting is dynamic and may be incomplete or delayed. it is also possible for the data owners to revise historical data. we based our assessment on data extracted in may 2020, but if this work is reproduced with data extracted at a later date, there may be some differences in the results. when using jupiter water-quality data for research purposes, it is imperative to implement stringent data-cleaning protocols. we set various quality control criteria (see supplementary file s1), resulting in the exclusion of pesticide analyses that did not meet those criteria. consequently, if raw jupiter data are used directly, or if different quality criteria are set, the results may differ from this assessment and our exposure estimates. the pesticides dataset is inherently heterogeneous and complex to work with due to the dynamic analytical programme, variable sampling frequency, high number of censored values and the variability around the lod. the dynamic analytical programme is a result of the frequent updates in national legislation in denmark regarding the list of mandatory pesticides for compliance monitoring. the final dataset structure reflects these changes in national legislation, but also the highly decentralised and asynchronous local response to newly discovered threats in groundwater catchments. currently, the list for mandatory compliance testing contains 46 pesticides (bek nr 1070, 28 october 2019), but new pesticides were included to the list in 2012 (n = 18), 2014 (n = 3); 2017 (dpc and methyl desphenyl chloridazon; mdpc), 2018 (1,2,4-triazol and dms) and 2019 (chlorothalonil sulfonic acid) (thorling et al. 2019). further, eight pesticides were removed in 2012, and https://doi.org/10.34194/geusb.v47.6090� www.geusbulletin.org� voutchkova et al. 2021: geus bulletin 47. 6090. https://doi.org/10.34194/geusb.v47.6090 13 of 16 www.geusbul let in.org 2,4-d was removed in 2015 (thorling et al. 2019). this changes the temporal or spatial data-coverage of individual pesticides (fig. 3), which should be considered in future uses of our final dataset. the different sampling frequency and consequently the varying length of data-gaps (fig. 4) can be explained by the legal requirements for minimum frequency of compliance testing, linked to the produced volumes of treated drinking water. for public waterworks, the mandatory sampling frequency varies from once every three years to once a month. however, water supply companies may choose to sample more often. as a result, the timeseries for individual pesticides and waterworks are highly irregular. combined with the large number of non-detects and the instability around the lod (fig. 5), interpolation to fill in missing concentration values in the pesticides timeseries should be avoided. we have introduced the qualitative group ‘not-analysed’ (or ‘unknown due to no data’) to retain this missing information. any statistical treatment of datasets with so many censored values poses a major challenge (helsel 2005; antweiler & taylor 2008; antweiler 2015; stow et al. 2018). the issue has been discussed specifically in the context of performing descriptive statistics and trend analysis of pesticides in danish groundwater (hansen et al. 2015, 2016; johnsen et al. 2016). due to the high percent of censored data in this dataset, simple substitution methods should be avoided. appropriate methods for the treatment of censored values are outlined in the references cited above. our assessment only required counting of analyses within the selected qualitative classes, and thus, substitution of concentrations <lod or loq was unnecessary. for the ‘pesticides-total’ (supplementary file s1), only the detected and quantified pesticides were used. 5.4.2 other input sources this assessment covers only the public waterworks with known locations and wsas where at least one pesticide was analysed within the study period. throughout the workflow, we excluded all pesticide analyses belonging to 95 public waterworks in denmark, of which 27 could not be geocoded (i.e. no coordinates or address) and 68 had no known wsa. the majority of the first group were supply network locations, which were registered incorrectly as separate waterworks. changes in the water supply structure could explain why some waterworks had no known wsa – all new waterworks that have become operational after 2014 would not be in the wsa dataset by schullehner and hansen (2014). additionally, changes may have occurred in the supply network of some of the waterworks, but evaluating such changes is beyond the scope of this article. an updated version of the wsa dataset is currently under preparation. some of the households included in this study may use private wells for potable water. approximately 3% of danish households use a private well as their source of drinking water (schullehner et al. 2017). we did not evaluate exposure to pesticides from private wells, since monitoring for drinking-water quality, especially for pesticides, is very limited for these private sources of water supply. however, private wells are known to often have significantly worse water quality than public supplies. in a regional study, carried out in 2000, 35% of 628 sampled private wells in denmark had pesticide concentrations >0.1 µg/l (brüsch et al. 2004). additionally, 22% of these wells violated the dwqs for nitrate (50 mg/l) and 48% violated the dwqs for microbiological contamination (brüsch et al. 2004). in a nationwide danish study, it was estimated that up to 30% of private wells violated the drinking-water standard for nitrate in the period 1978–2012 (schullehner et al. 2017). therefore, we are likely underestimating the percent of households where pesticide levels exceed the dwqs. our status assessment and exposure estimation were performed for four general qualitative classes because of limitations in the pesticide dataset. care must be taken not to over interpret results. for the status assessment at waterworks, one analysis for an individual pesticide exceeding the dwqs within the selected periods would classify the waterworks as ‘exceeding dwqs’. similarly, if one waterworks within a wsa had at least one exceedance in the respective period, all households in that area would be classified as exposed to concentrations exceeding the dwqs. therefore, we are likely overestimating some of the exposure for wsas with multiple waterworks. our assumptions correspond to a worst-case scenario which is in line with deterministic approaches for estimating exposure to pesticides (ferrier et al. 2002). to quantify this potential overestimation, we categorise the households potentially exposed to >0.1 µg/l in three groups based on the percentage of waterworks that exceed the dwqs in the wsa (fig. 8; table sm-5 in supplementary file s1). in the first group, where 100% of the waterworks in a wsa exceeded dwqs, the percent of households potentially exposed to pesticides >0.1 µg/l is 5.1% (2002–2019) and 4.5% (2015–2019). in the second group, where ≥50% of the waterworks in a wsa exceeded dwqs, these percentages are 8.7% (2002–2019) and 7.1% (2015–2019). the third group is the most uncertain group, where <50% of the waterworks in a wsa exceeded the dwqs. this corresponds to 10.2% (2002–2019) and 4.2% (2015–2019) of households. in most of these wsas, however, the other waterworks contained detected pesticides (0.03–0.10 µg/l). in our study, we were not able to calculate exposure with higher spatial resolution without knowing more about the distribution patterns in the supply system within the wsa. https://doi.org/10.34194/geusb.v47.6090� www.geusbulletin.org� voutchkova et al. 2021: geus bulletin 47. 6090. https://doi.org/10.34194/geusb.v47.6090 14 of 16 www.geusbul let in.org our assessment provides a first overview on the state of danish treated drinking water and the potential exposure at household level based on the selected qualitative exposure classes. more detailed assessments on the level of individual consumers can be achieved by epidemiological studies exploring the potential association between exposure to pesticides from drinking water and various health outcomes. 6 conclusions and perspectives our study provides a national assessment of pesticide status of danish drinking water supplied by public waterworks and an estimation of household exposure to pesticides from drinking water. the national groundwater assessment and the danish epa mass screening have already provided evidence that pesticide pollution of danish groundwater is widespread and not only an issue of local importance. our results compliment these findings and quantify the status based on the treated drinking water supplied to the danish population. we found that: 1. 0.5% (n = 4925) of the individual analyses of treated drinking water contained a quantified pesticide (≥0.03 µg/l) and of those 16% (n = 793) exceeded the dwqs. 2. bam had the most complete spatiotemporal coverage, and dpc had the highest measured concentration. together with dms, these three pesticides had the highest number of dwqs exceedances in the study period. 3. 9% (2002–2019) and 7% (2015–2019) of the waterworks had an exceedance of the dwqs (>0.1 µg/l). in addition, 20% and 17% of waterworks contained pesticides with lower concentrations (0.03–0.10 µg/l) for the two respective periods. these results compare well with the raw groundwater status reported by thorling et al. (2019). 4. 19% (2002–2019) and 11% (2015–2019) of danish households were potentially exposed to pesticides >0.1 µg/l. in addition, 37 and 30% were potentially exposed to lower concentrations (0.03–0.10 µg/l) for the two periods, respectively. 5. one-third of the pesticides with established who guideline values (protective against health effects from a lifetime exposure) were found, but all were lower than the guideline values. we demonstrate that the central registration of all chemical analyses of danish drinking water in the jupiter database allows us to assess the current and past spatiotemporal status of danish drinking water. while the open access to jupiter is a major advantage, care should be taken as the data structure is highly heterogeneous regarding sampling frequency and pesticides analysed. a strength of our study is that we implemented and documented extensive data pre-processing procedures in detail, allowing reproducibility and further informed use of the dataset. we have also included a comprehensive account of different methodological limitations, which are important for future epidemiological studies. acknowledgements we thank ingelise m. balling and lærke thorling for providing comments and suggestions on the data quality control and filtering procedure (qc1). we are also thankful to lærke thorling and anders r. johnsen, whose thoughtful comments on an earlier version of this manuscript helped us clarify and sharpen some of the findings. we are grateful to the two anonymous reviewers for providing constructive and detailed reviews, which pushed us to perform additional uncertainty analysis and to restructure the manuscript for clarity. additional information funding statement this work was funded under the project ‘geographical clustering of leukaemia and multiple myeloma and association with pesticides in ground water’ by the karen elise jensen fond (2017–2021). author contributions dv contributed to the conceptualisation, data curation, formal analysis, methodology, visualisation and writing (original draft). js contributed to the conceptualisation, data curation, methodology, validation and 2002–2019 5.1 3.6 10.2 4.5 2.6 4.22015–2019 19% 11% 0% 5% 10% 15% households (%) potentially exposed to pesticides >0.1 µg/l by wsa type 100% of the waterworks had at least once pesticide exceeded 0.1 µg/l ≥50% of the waterworks had at least once pesticide exceeded 0.1 µg/l <50% of the waterworks had at least once pesticide exceeded 0.1 µg/l water supply areas (wsas), where: fig. 8 uncertainty in the estimate of households (%) exposed to pesticides exceeding the dwqs (0.1 µg/l). https://doi.org/10.34194/geusb.v47.6090� www.geusbulletin.org� voutchkova et al. 2021: geus bulletin 47. 6090. https://doi.org/10.34194/geusb.v47.6090 15 of 16 www.geusbul let in.org writing (review and editing). cs contributed to the conceptualisation, validation and writing (review and editing). kw contributed to the conceptualisation, validation and 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https://www.eionet.europa.eu/etcs/etc-icm/products/etc-icm-report-1-2020-pesticides-in-european-rivers-lakes-and-groundwaters-data-assessment� https://www.eionet.europa.eu/etcs/etc-icm/products/etc-icm-report-1-2020-pesticides-in-european-rivers-lakes-and-groundwaters-data-assessment� https://naturstyrelsen.dk/media/nst/attachments/videregende_vandbehandling_2012.pdf� https://naturstyrelsen.dk/media/nst/attachments/videregende_vandbehandling_2012.pdf� https://cirrau.au.dk/fileadmin/cirrau/documents/ophold2016b.pdf� https://cirrau.au.dk/fileadmin/cirrau/documents/ophold2016b.pdf� http://qgis.osgeo.org� http://qgis.osgeo.org� https://www.r-project.org/� voutchkova et al. 2021: geus bulletin 47. 6090. https://doi.org/10.34194/geusb.v47.6090 16 of 16 www.geusbul let in.org rosenstock, l. 1991: chronic central nervous system effects of acute organophosphate pesticide intoxication. the lancet 338, 223–227. https://doi.org/10.1016/0140-6736(91)90356-t rstudio team. 2018: rstudio: integrated 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henry, l. 2019: tidyr: tidy messy data. r package version 1.0.0. https://cran.r-project.org/package=tidyr world health organization. 2011: guidelines for drinking-water quality, 4th edition. geneva: world health organization. https://www.who.int/ publications/i/item/9789241548151 (accessed january 2020) world health organization. 2017: guidelines for drinking-water quality, 4th edition, incorporating the first addendum. geneva: world health organization. https://www.who.int/publications/i/item/9789241549950 (accessed july 2020) world health organization. 2019a. preventing disease through healthy environments. exposure to highly hazardous pesticides: major public health concern. geneva: world health organization. https://apps.who. int/iris/handle/10665/329501 (accessed july 2020) world health organization. 2019b: who recommended classification of pesticides by hazard and guidelines to classification, revision 2019. geneva: world health organization. https://apps.who.int/iris/handle/10665/329501 (accessed july 2020) https://doi.org/10.34194/geusb.v47.6090� www.geusbulletin.org� https://doi.org/10.1016/0140-6736(91)90356-t� http://www.rstudio.com/� https://doi.org/10.1088/1748-9326/9/9/095001� https://doi.org/10.1016/j.gexplo.2017.03.006� https://doi.org/10.1016/j.chemosphere.2019.06.207� https://doi.org/10.1073/pnas.1500232112� https://doi.org/10.1016/j.ecoinf.2018.03.002� https://doi.org/10.1016/j.ecoinf.2018.03.002� https://doi.org/10.3390/geosciences5010045� https://ggplot2.tidyverse.org� https://cran.r-project.org/package=dplyr� https://cran.r-project.org/package=dplyr� https://cran.r-project.org/package=tidyr� https://www.who.int/publications/i/item/9789241548151� https://www.who.int/publications/i/item/9789241548151� https://www.who.int/publications/i/item/9789241549950� https://apps.who.int/iris/handle/10665/329501� https://apps.who.int/iris/handle/10665/329501� https://apps.who.int/iris/handle/10665/329501� https://apps.who.int/iris/handle/10665/329501� estimating pesticides in public drinking water at the household level in denmark abstract 1 introduction 1.1 pesticide exposure and drinking water 1.2 objectives 2 study setting and legal framework 3 methods and materials 3.1 data sources 3.1.1 drinking water samples analysed for pesticides 3.1.2 geographic coordinates of the waterworks 3.1.3 water supply areas of the waterworks 3.1.4 per cent of households located within a wsa 3.2 data joining procedure 3.3 final pesticide dataset 3.3.1 limit of detection and quantification 3.3.2 structural heterogeneity 3.4 status assessment and exposure estimation 3.5 software 4 results 4.1 status overview 1 individual compounds 4.2 status at waterworks and wsas (status 2 and 3) and pesticide exposure 5 discussion 5.1 possible trends 5.2 comparison with untreated groundwater 5.3 drinking water quality standard and health implications 5.4 limitations 5.4.1 drinking water pesticide data 5.4.2 other input sources 6 conclusions and perspectives acknowledgements additional information references figures fig. 1 workflow including data-processing steps and overview of excluded chemical analyses and water fig. 2 number of public waterworks with pesticide data for each water supply area (wsa) for a: the e fig. 3 proportion of waterworks with ‘unstable’ detections of pesticides during the study period (20 fig. 4 data availability. a: period (in years) where data are available for individual pesticides, c fig. 5 frequency of sampling events by sampling date in a: 2002-2019 and b: 2015-2019. the cumulativ fig. 6 pesticide status at waterworks and estimated exposure at the household level. note: the class fig. 7. pesticide status at waterworks and water supply area (wsa). ‘never detected’ refers to all a fig. 8 uncertainty in the estimate of households (%) exposed to pesticides exceeding the dwqs (0.1 µ tables table 1 results from the pesticide status of untreated groundwater (thorling et al. 2019) table 2 drinking water guideline values for pesticides (world health organization 2017) and occurren untitled 33 lithostratigraphy the upper pleistocene in the vendsyssel region comprises three major stratigraphic units: (1) a unit of marine sediments deposited on the erosional surface of the saalian till, overlain by (2) a glacioterrestrial succession that in turn is succeeded by (3) a second marine deposit. the first marine unit was laid down after the retreat of the ice from the region at the end of saalian time. during eemian and early to middle weichselian time, the older yoldia sea prevailed (jessen et al. 1910; jessen 1918). from the end of the middle weichselain to the latest late weichselian, the area was subjected to terrestrial glaciation (houmarknielsen et al. 1996). after the ice melted back from the main stationary line, a marine environment was re-established and persisted until isostatic rebound resulted in subaerial exposure of the seabed of the younger yoldia sea (jessen 1918; figs 12, 13). a new lithostratigraphical subdivision is proposed to cover the three upper pleistocene successions (fig. 14). the systematic stratigraphic framework is based on formations defined according to the guidelines given by rawson et al. (2002). the lower marine unit, corresponding to the deposits representing the older yoldia sea and formerly referred to as the skærumhede series (jessen et al. 1910), is here defined as the skærumhede group (new group). the group includes the middle weichselian stortorn and lønstrup klint formations (new formations) and an unnamed lower unit mainly including the eemian and lower weichselian deposits (figs 15, 16). four formations are distinguished in the glacioterrestrial unit: the glaciofluvial and glaciolacustrine rubjerg knude formation (new formation), the kattegat till formation (houmark-nielsen 1987, 1999, 2003), the ribjerg formation (new formation) and the mid danish till formation (houmark-nielsen 1987, 1999, 2003). the uppermost major unit, comprising the post-glacial arctic marine younger yoldia clay and saxicava sand of jessen (1918, 1931), is referred to the vendsyssel formation (new formation) (fig. 14). skærumhede group new group history. the skærumhede group includes most of the lithological units formerly described as the skærumhede series (jessen et al. 1910). these include the marine eemian, the marine lower weichselian and the marine–brackish–lacustrine beds in the middle weichselian (figs 14, 15; lykke-andersen & knudsen 1991; knudsen 1994). recognition of the group is primarily based on a research borehole behind the farm at skærumhede, about 10 km west of frederikshavn (fig. 13), that was drilled by the geological survey of denmark to investigate the source of natural gas in the vicinity of frederikshavn (jessen et al. 1910). the well penetrated to a depth of 235 m and terminated in upper cretaceous chalk. above the chalk, a 20 m thick unit of till and glacial sediments was encountered. the till is now referred to the saalian (lykke-andersen 1987), and forms the basal unit of the quaternary succession over most of north jylland (fredericia 1982, 1983a, b; pedersen 1989). the succession above the saalian glacial sediments was described under the heading: ‘the marine skærumhede series’ by jessen et al. (1910 pp. 67, 156). this unit is c. 123 m thick, from 57.4 m to 180.3 m below surface, corresponding to a lower boundary at 157.1 m and a top at 34.2 m below sea level. it was subdivided into three biostratigraphic zones: 1) the turritella terebra zone (74 m thick), 2) the abra nitida zone (8.5 m thick) and 3) the portlandia arctica zone (40 m thick) (jessen et al. 1910). additional details were added to the unit based on several glaciotectonically dislocated outcrops in the northern part of vendsyssel by jessen et al. (1910) and jessen (1918, 1931). subsequent discussion concerning the stratigraphic position and dif ferentiation of the skærumhede series resulted in a new borehole, which was directed by the geological survey of denmark at the skærumhede locality in the early 1970s. although the borehole only went down to 120 m below surface, it gave a good record of the lithology and macrofauna and in particular provided samples for a detailed foraminiferal investigation (bahnson et al. 1974). facing page: fig. 14. schematic stratigraphic log of the units represented in the rubjerg knude glaciotectonic complex. the fossils indicated on the log represent 14c-dated samples. 34 name. the skærumhede group is named after the locality of skærumhede c. 10 km west of frederikshavn, denmark (fig. 13). type section. the type section is defined as the skærumhede well (dgu no. 10.4 and 10.392) (fig. 15), where the pioneer drill site for natural gas was situated at a barren and unfertile place caused by seepage of gas from the subsurface (fig. 13; jessen et al. 1910). reference sections. reference sections are proposed in well-documented borehole sections: the nørre lyngby ii well (dgu no. 8.137) described by lykke-andersen (1987), and the skagen iii well (dgu no. 1.287) recorded by knudsen (1994) and petersen (2004) (fig. 16). lithology. the skærumhede group consists of rather uniform bluish-black to dark grey clay with minor intercalations of silt and fine-grained sand. the silt laminae and thin fine-grained sand beds become more common towards the top of the group. macrofossils vendsyssel formation rubjerg knude formation lønstrup klint formation stortorn formation skærumhede group (undiff.) clay silty mud dropstones in mud sandy mud sand till gravel chalk sk æ ru m he de g ro up 10 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 210 220 230 clay silt sand cobble pebblef m c skærumhede well w ei ch se lia n u . c re ta ce ou s sa al ia n ee m ia n m below surface (e. 23 m a.s.l.) marine fossil fig. 15. lithostratigraphic log of the skærumhede well dgu no. 10.4, the type section of the skærumhede group and the stortorn formation. 35 are present through most of the group but decrease in abundance towards the top (jessen et al.1910). dropstones are present in the middle of the group and increase in abundance towards the uppermost part, in which graded silts and sands are intercalated with grey mud. boundaries. the lower boundary is the unconformity on top of the saalian till. the upper boundary is an unconformity overlain by coarse clastic sediments interpreted as a residual boulder bed (unit b of sadolin et al. 1997). thickness. the thickness of the group varies from nearly 130 m in the type section in the skærumhede well, to c. 48 m in the nørre lyngby well and c. 45 m in the skagen iii well (fig. 16). distribution. knowledge of the distribution of the group in the vicinity of rubjerg knude is based on the skærumhede well (dgu no. 10.4), the nørre lyngby well (dgu no. 8.137) and the skagen well (dgu no. 1.287; figs 13, 16). the group is also known from wells in the northern part of vendsyssel and the islands of læsø and anholt in the kattegat. according to these records, the group extends from the western part of the kattegat at frederikshavn and læsø, towards the south central part of vendsyssel, from where it continues offshore between anholt and djursland (knudsen 1994, fig. 3). the group extends offshore into the strait between læsø and the swedish coast. the southward extent is not known, but is probably up to about 30 km south of anholt. the extent to the north is also uncertain and has not yet been mapped. it is inferred that it may occur in the western part of the skagerrak (knudsen 1994) and it may also extend out into the northern part of the north sea. age. the age of the group extends from the beginning of the eemian, about 130 000 years b.p. (knudsen 1994), to the latest part of the middle weichselian, about 30 000 years b.p. (houmark-nielsen 1999). depositional environment. at the lower boundary of the group, red corroded flints were recognised in the skærumhede well (jessen et al.1910) indicating that the top surface of the saalian till had been exposed and subjected to subaerial erosion prior to the transgression that culminated in the eemian. during the eemian, a deep-water shelf environment was established with water depths exceeding 100 m; in the early weichselian, water depths decreased dramatically to less than 50 m (knudsen 1994). the decrease in water depth continued during the middle weichselian under increasing glacial influence. subdivisions. the upper skærumhede group is subdivided into the stortorn and lønstrup klint formations. the lower part of the group, mainly including the marine eemian and lower weichselian deposits, is presently undifferentiated. stortorn formation new formation history. in the lønstrup klint section, two units of grey-bluish clay subjected to glacial deformation have been distinguished, the diluvial clay and the portlandia arctica clay (jessen 1931). the latter unit corresponds to the so-called older yoldia clay (ældre yoldialer in danish), which in the skærumhede well was referred to as the portlandia arctica zone and in the crosssection of lønstrup klint is indicated to occur at three localities (jessen 1931). the most impressive of these is the stortorn site, where dark grey – black clay, rich in mollusc shells, crops out (fig. 21). the site is inaccessible, or difficult of access, since the slippery clays occur in the breaker zone at the foot of the almost vertical cliff section. the other two localities are the cliff sections just beyond the town lønstrup, locally named ‘lille blå’ (little blue), and the clif f section below the northern corner of the mårup churchyard. at all three sites, the unit is tectonically disturbed which hampers detailed logging of the succession. in addition, the formation occurs locally in the lower thrustsheet duplexes north of mårup church and in the moserende cliff section (see description of that section, below). name. the formation is named after the stortorn cliff section at lønstrup klint. the formation is here incorporated within the lowermost thrust unit in the stortorn section (see below). type section. the type section for the formation is the skærumhede well, dgu no. 10.4 (figs 14, 15). reference sections. the reference sections for the formation are the outcrops at stortorn and north of the northern corner of mårup churchyard (fig. 17) in the 36 weichselian saalian holocene upper cretaceous eemian vendsyssel fm kattegat till fm rubjerg kn. fm lønstrup kl. fm lønstrup kl. fm stortorn fm stortorn fm skærumhede group (undiff.) stortorn fm lønstrup klint fm clay silt sand cobble pebblef m c 10 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 10 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 210 220 230 clay silt sand cobble pebblef m c silt clay sand cobble pebblef m c nørre lyngby m below surface (e. 20 m a.s.l.) skærumhede m below surface (e. 23 m a.s.l.) skagen m below surface (e. 3 m a.s.l.) 10 20 30 40 50 60 70 80 90 100 110 120 130 140 150 v endsyssel fm clay silty mud dropstones in clay and silty mud sandy mud sand till gravel chalk marine fossils vend sys sel fm rubj erg kn. f m 37 mud sand f. m. c. gr.pb.co. 0 m 5 10 15 20 25 sandy mud sand dropstones in mud gravel shells lamination structureless climbing ripples trough cross-bedding ball-and-pillow / convolute bedding thrust fault clayey mud current ripple cross-lamination vendsyssel formation dune sand lønstrup klint formation sk æ ru m he de g ro up x aar 4069 stortorn formation unconformity thrust zone fig. 17. sedimentological log of the succession in the southern part of the ribjerg section (above the ‘store blå’). the stortorn formation records an arctic marine deposit, yielding shells typical of this environment: hiatella arctica, mya truncata and portlandia arctica. the boundary between the stortorn and lønstrup klint formations constitutes a thrust-fault breccia indicating dif ferentiation into thrust-fault duplex segments of the skærumhede group. the vendsyssel formation at the top of the section was deposited on an erosional unconformity with a lag conglomerate at the base. the location of the sample collected for 14c dating (aar 4069) is indicated (see table 2). facing page: fig. 16. simplified lithological logs from three thoroughly documented wells in vendsyssel (nørre lyngby: dgu no. 8.137; skærumhede: dgu no. 10.4; skagen: dgu no. 1.287, for location see fig. 13). the logs illustrate the stratigraphic correlation of the units defined in the investigation of the rubjerg knude glaciotectonic complex. the difference in thickness of eemian–weichselian deposits mainly reflects the average content of sand; the skagen well represents a deeper marine depositional environment compared to the skærumhede well. note that the cretaceous deposits in the skagen well comprise turonian–cenomanian greensands. the figure is based on information from jessen et al. (1910), bahnson et al. (1974), lykke-andersen (1987), lykke-andersen & knudsen (1991), knudsen (1994) and petersen (2004). 38 lønstrup klint coastal cliff. the nørre lyngby well (dgu no 8.137, lykke-andersen 1987) is the well closest to stortorn where the undisturbed formation has been penetrated. additional sections include the coastal clif f at hirtshals displaying allochthonous peats in the black clay formation, and the skagen iii well (dgu no 1.287; petersen 2004) that includes a clay unit, 8 m thick, here referred to the stortorn formation (fig. 16). lithology. the stortorn formation consists of black, locally dark grey – bluish structureless clay with a large number of dropstones, which are commonly glacially striated. lenses or irregular beds, up to 10 cm thick, of shell debris (gravel-size) occur scattered in the unit, and the abundance of shells in local patches gives the formation a white spotted appearance (fig. 18). at the top of the formation, the clayey mud changes colour from dark bluish grey to violet-brown and develops recognisable lamination. fossils. the unit has been referred to the portlandia arctica zone of jessen et al. (1910) since this is the most abundant mollusc species in the clay (bahnson et al. 1974). macoma calcarea is another common mollusc and hiatella arctica occurs in abundance. a list of characteristic molluscs and their distribution in the unit is given by k.s. petersen (in: bahnson et al. 1974). moreover, the presence of balanus sp. and additional erratic macrofossils are reported. the most common microfossils are the foraminifers elphidium excavatum and cassidulina crassa. p.b. konradi and k.l. knudsen (in: bahnson et al . 1974) documented and discussed the foraminiferal fauna. boundaries. the lower boundary of the formation is defined by a shift from clayey mud to mud with a marked increase in coarse-grained ice-rafted debris. the increased content of coarse-grained material is associated with an abundance of mollusc shells and fragments. the upper boundary of the formation is defined at the transition from marine clay showing diffuse lamination and colours varying from grey bluegreen to violet-brown, to a grey clayey and silty mud intercalated with graded silt and fine-grained sand laminae a few millimetres thick. thickness. the formation is about 20 m thick. in the skagen iii well, the formation is only about 8 m thick, probably due to the more offshore position and deeper water environment in this part of the basin (petersen 2004). distribution. the distribution of the stortorn formation is identical with the distribution of the skærumhede group. the formation can be readily identified in the nørre lyngby well (lykke-andersen 1987; figs 13, 16) and it has also been described from the hirtshals cliff section (lykke-andersen 1971). in addition, it is known from the deeper wells in the main part of the vendsyssel area and from the islands of læsø and anhiatella shells 5 cm fig. 18. the stortorn formation in the stortorn section is dominated by black sticky clay. locally, shells of hiatella arctica and portlandia arctica are very abundant. dropstones are also common in the formation. photograph: august 2001. 39 holt (fredericia 1982, 1983a, b, 1984; lykke-andersen 1987). age. three shell samples from the stortorn formation at lønstrup klint have been 14c dated, using the atomic mass spectrometric (ams) method, for the present investigation. two of the samples were derived from the archives of the former geological survey of denmark; two shells of hiatella arctica were chosen for dating the formation at the stortorn locality and from the northernmost outcrop of the formation at lønstrup klint (‘lille blå’, at the base of the northern part of the ribjerg section, collected and described by a. jessen) (table 2). the third sample was taken in 1996 and comprises shells of hiatella arctica from the shellbearing clay outcrop at stortorn (reference dgu no. 00136, aar-4069, table 2). these were dated to test the collection made nearly 100 years earlier and provided an age for the lowermost stortorn formation, namely 31 300 (± 400) years b.p. the age of the upper levels of the formation, as represented by the muddy sediments at the ‘lille blå’ section is slightly younger (30 000 (± 400) years b.p. (table 2)). the new dating of the stortorn formation corresponds well with previous age dates from the upper part of the skærumhede group, which gave 32 000 years b.p. (seidenkrantz & knudsen 1993). depositional environment. the presence of a boreal fauna including mytilus edulis, arctica islandica and zirphaea crispata in the shell-debris gravel in an environment characterised by a bottom fauna of portlandia arctica and macoma calcaria led nordmann (1928) and jessen (1931) to conclude that the boreal shallow-water faunas of interglacial affinity were transported as ice-rafted material into more offshore arctic marine environments. the most convincing examples of such erratic material are the dropstones with balanus sp. the stortorn formation is thus interpreted to have been deposited during a period of decreasing water depths in a marine environment characterised by dispersal of erratics from drifting icebergs. lønstrup klint formation new formation history. the characteristic development of this formation, viz. grey clayey muds interbedded with layers of fine-grained sand, occurs in the steeply inclined sheets that are prominent in the cliff of lønstrup klint. due to the absence of macrofossils, this unit was named diluvialler (diluvial clay) and the unit was correlated with the uppermost part of the skærumhede group in the skærumhede well (jessen et al. 1910; jessen 1918, 1931). the sedimentology of the formation was described from sandrende at lønstrup klint under the heading unit a by sadolin et al. (1997). name. the formation is named after the coastal cliff of lønstrup klint. type section. the type section for the formation is at sandrende in lønstrup klint, situated between point 3500 and 3600 m in the rubjerg knude cross-section (plate 1), from where a sedimentological log was provided by sadolin et al. (1997) (figs 13, 19). reference sections. reference sections are defined at ulstrup rende (fig. 20) and at kramrende (fig. 21) situated at 5950 m and 4500 m, respectively, in the rubjerg knude cross-section (plate 1). moreover, the skærumhede and the skagen wells are reference sections for the western and northern development of the formation (fig. 16). lithology. the formation consists of blue-grey clayey and silty laminated mud and cross-laminated beds of fine sand. the lowest part of the formation is characterised by dark grey mud, interlayered with laminated to thin-bedded clayey and silty mud. the light grey beds, 1–5 cm thick, grade upwards from light grey silt to dark grey clay (fig. 22). some of the dark grey clayey mud levels are interbedded with thin lenticular, light-coloured silt and fine-grained sand laminae (fig. 23). silty to fine-grained sandy beds may be up to 1 m thick. dropstones occur scattered in the bluegrey mud. in the upper part of the formation, beds of light grey sand, 3–8 m thick, occur interbedded with a few thin beds of laminated mud. the thick sand beds are characterised by climbing ripple cross-lamination. in the cliff section at rubjerg knude, much of the primary bedding in the lønstrup klint formation is disturbed by water-escape structures (ball-and-pillow etc.) and hydrodynamic brecciation (flame to diapir structures). fossils. macrofossils have not been found in the formation and the foraminifers, dominantly elphidium excavatatum, are interpreted to be redeposited (lykkeandersen 1987). 40 mud sand f. m. c. gr.pb.co. 5 10 15 0 m l/r-unconformity 20 25 30 35 40 sandy mud sand gravel basal till lamination structureless climbing ripples trough cross-bedding slump structures ball-and-pillow / convolute bedding thrust fault rubjerg knude formation lønstrup klint formation kattegat till formation clayey mud current ripple cross-lamination fig. 19. sedimentological log of the succession in the sandrende section, including the type sections of the lønstrup klint and rubjerg knude formations. the lønstrup klint formation represents lacustrine deposition, whereas the rubjerg knude formation records a shift from fluvial to lacustrine sedimentation, returning to fluvial sedimentation in the upper levels (sadolin et al. 1997). note the synsedimentary small-scale thrust structures that appear in the upper levels of the rubjerg knude formation indicating that the formation was deposited in a piggyback basin. 41 boundaries. the lower boundary of the lønstrup klint formation is placed at the transition from the typical marine clay with a macrofossil fauna of the stortorn formation to unfossiliferous laminated clayey and silty muds with laminae and thin beds of fine sand. the upper boundary is the marked erosional unconformity between the lønstrup klint and rubjerg knude formations, referred to as the l/r-unconformity (figs 14, 24). thickness. the maximum thickness of the formation is about 25 m, but with large variations due to erosional relief at the l/r-unconformity. distribution. the lønstrup klint formation is distributed over the main part of vendsyssel. it is erosionally truncated at the top towards the south where underlying glacial deposits or upper cretaceous chalk constitute the surface geology of the coastal areas north of limfjorden. to the east, it probably extends offshore into the middle part of kattegat from where it is known in wells on the islands of læsø and anholt (lykke-andersen 1987; knudsen 1994). the extent out into the north sea to the west remains unknown. towards the north, it extends offshore into the skagerrak beyond skagen, where it is recorded in the skagen iii well (petersen 2004). 24 20 15 10 5 0 m mud sand f. m. c. gr. pb.co. glacitectonite rubjerg knude formation ulstrup glaciofluvial beds lønstrup klint formation clayey mud sandy mud sand gravel lamination current ripple cross-lamination trough cross-bedding tectonite water-escape pipes and sand-filled cracks structureless sediment l/r-unconformity fig. 20. sedimentological log of the succession in the southernmost thrust sheet in the ulstrup section. the log was measured at ulstrup rende, situated at point 5980 m in the cross-section (plate 1). note the 4 m thick sand-crack dominated thrust zone that characterises the lower part of the lønstrup klint formation, which developed during thrust-sheet translation along the hangingwall flat (figs 46, 47). note also that in the rubjerg knude formation, the lowermost 7 m corresponds to the glaciofluvial ‘ulstrup beds’ (figs 28, 29). 42 age. 14c dating of plant debris from the formation indicates an age of about 30 000 b.p. (houmark-nielsen et al. 1996). this age is compatible with the 32 000 b.p. age derived from the underlying stortorn formation. depositional environment. the transition from the stortorn formation to the non-fossiliferous lønstrup klint formation is interpreted as a shift from a normal, arctic marine environment through brackish to a freshwater environment dominated by rapid deposition of suspended sediment supplied to the basin by meltwater. the sharp-based normally graded silt and sand beds are interpreted as fine-grained turbidites. the dark mud was deposited from suspension, whereas the lenticular sand/silt laminae represent wave-reworked, distal storm-sand layers. sedimentation started below storm-wave base and it is suggested that the lake environment was deep and of fairly wide extent. the occurrence of numerous fine-grained sandy turbidites sourced mainly from the south probably reflects exposed land areas in the southern part of vendsyssel during the low stand of sea level (sadolin et al. 1997). the sand beds are interpreted to record relatively rapid sedimentation by sediment gravity 0 m 5 10 15 20 25 thrust zone rubjerg knude formation lønstrup klint formation sandy mud sand gravel lamination structureless climbing ripples trough cross-bedding ball-and-pillow / convolute bedding tectonite clayey mud current ripple cross-lamination mud sand f. m. c. gr.pb. co. l/r-unconformity fig. 21. sedimentological log of the lønstrup klint and rubjerg knude formations in the kr01 thrust sheet in the southern part of the kramrende section (point 4500 in plate 1). the fine-grained, thinto medium-bedded sandy turbidites in the lønstrup klint formation are interbedded with thin layers of blue-grey silty mud. these sand beds are often disrupted into ball-and-pillow load structures (see fig. 48). note the tectonite at the base of the succession, related to the hanging-wall flat of the kr01 thrust sheet. 43 flows in a glaciolacustrine environment; the thickest of these beds may represent deposition within one summer of sediment derived from the southern slopes of the basin (sadolin et al.1997). hydrodynamic deformation of the strata was initiated at a very early stage in the glaciotectonic process, as loading by the superposed rubjerg knude formation and by glaciotectonic thrust sheets resulted in increasing pore-water pressure. hydrodynamic brecciation continued during deformation until the displacements of thrust sheets ceased. upper weichselian lithostratigraphic units rubjerg knude formation new formation history. in the steep cliff section at rubjerg knude, the thrust sheets, consisting of the grey-blue coloured clay of the lønstrup klint formation, are depositionally overlain and structurally underlain by light-coloured yellowish sand, named diluvialsand (diluvial sand) by jessen (1918, 1931). the succession was refig. 22. laminated to thin-bedded clayey and silty mud in the lower part of the lønstrup klint formation in the rubjerg knude fyr section. the bedding is defined by layers grading from light grey silt to dark grey clay, and the sharp-based normally graded beds are interpreted as fine-grained turbidites. the coin for scale is 2.5 cm in diameter. photograph: september 1985. fig. 23. dark grey clayey mud interbedded with thin lenticular, light coloured silt and fine-grained sand laminae. the mud was deposited from suspension, whereas the lenticular laminae represent wave-reworked, distal storm-sand layers, deposited below storm-wave base. photograph: june 1993. 44 fig. 24. the l/r-unconformity is steeply inclined in the gr08 thrust sheet. note the large-scale cross-bedding in the basal unit of the rubjerg knude formation, which onlaps the unconformity (r-onlap). photograph: june 1993; rucksack for scale. l/r-unconformity 0 m 13 10 ulstrup glaciolacustrine beds lønstrup klint formation rubjerg knude formation thrust zone 5 mud sand f. m. c. gr.pb. co. sandy mud sand gravel lamination structureless sediment trough cross-bedding ball-and-pillow convolute bedding water-escape pipes and sand-filled cracks thrust fault clayey mud current ripple cross-lamination fig. 25. sedimentological log of the succession in the northern thrust sheet in the ulstrup section. the log was measured near the ulstrup steps at point 5625 m in the cross-section (plate 1). the base of the log is the hanging-wall flat of the ul02 thrust sheet and the lowermost 2 m constitute the thrust zone. the boundary between the lønstrup klint formation and the ulstrup glaciolacustrine beds above (lower unit of the rubjerg knude formation) is a flat, non-erosional surface but can be traced to the northern part of the ul02 thrust sheet where this boundary is a clear erosional unconformity. 45 ferred to as units b–d in the sedimentological study by sadolin et al. (1997). name. the formation is named after rubjerg knude, the highest part of the lønstrup klint cliff.. type section . the type section is at sandrende in the lønstrup klint cliff section (figs 13, 19). reference section. four reference sections are defined, all situated in the vicinity of rubjerg knude. in the distal part of the rubjerg knude glaciotectonic complex, two reference sections are defined at ulstrup. the first of these is located in ulstrup rende at point 5900 m in the rubjerg knude cross-section (plate 1) which demonstrates the presence of coarse-grained glaciofluvial channel fill deposits (fig. 19). the second reference section at ulstrup is located at point 5450 m in plate 1 and documents the occurrence of fine-grained clayey muddy glaciolacustrine beds in the formation (fig. 25). the third reference section is situated at martørv bakker (point 4850 m in plate 1), which demonstrates diamictitic sediments including slump units in a piggyback basin (fig. 26). the fourth reference section is located at moserende at point 1750 m in plate 1. this section illustrates the formation in a piggyback basin situated in a proximal position in the rubjerg knude glaciotectonic complex (fig. 27). 23 mud sand f. m. c. gr.pb. co. 20 15 10 5 0 m vendsyssel formation holocene peat ( martørv ) aeolian sand rubjerg knude formation (including diamict sediments and slump-folded units) lønstrup klint formation l/r-unconformity erosional unconformity sandy mud sand gravel lamination structureless sediment trough cross-bedding mud with scattered pebbles and cobbles slump structures ball-and-pillow convolute bedding clayey mud current ripple cross-lamination fig. 26. geological log of the diamict sediments and slump-fold structures, which represent the rubjerg knude formation in the piggyback basin on the back of the mb02 thrust sheet in the martørv bakker section, point 4880 in the cross-section (plate 1). 46 lithology. the dominant lithology of the formation is fineto medium-grained sand. beds of gravel occur in the lowermost 1–5 m, related to the initial deposition succeeding the formation of the erosional unconformity (the l/r-unconformity) (figs 24, 27, 28, 29). the sediment source was partly the main central part of the danish basin, indicated by the content of 23–25% flint and upper cretaceous chalk, and partly outwash material from the propagating ice margin, indicated by the c. 75% basement clasts (jessen 1931). many sand beds display small-scale current ripple lamination, and some show well-developed climbing ripples (sadolin et al. 1997). large-scale cross-bedded sand is observed in shallow channel fills, and some trough cross-stratification occurs in relation to growth-fault structures formed along normal faults or depressions related to the formation of synsedimentary footwall synclines (see fig. 87). a series of large-scale accretionary cross-stratification structures are related to a shift in the substratum inclination during thrust-fault propagation (see figs 82, 95). clasts of clay derived from the lønstrup klint formation are common, and in some of the syntectonic settings these beds rich in clay-clasts may be regarded as sedimentary clastic breccias with olistoliths or lumps of sandy mud 1–5 m in size (fig. 56). the olistoliths represent the frontal parts of thrust sheets, which gravity-glided out into depressions formed during thrust-fault propagation. locally, some of the depressions developed into small glaciolacustrine basins characterised by interbedded finegrained sands and sandy muds with current crosslamination; these deposits may reach a thickness of up to 5 m (fig. 25). fossils. redeposited fossils occur together with accumulations of twigs and amber (‘ravpindelag’). wellpreserved arctic mosses suitable for 14c dating the formation have been separated from the organic debris (houmark-nielsen et al 1996). in the basin at stensnæs, a large number of the mollusc shells were rel/r-unconformity mud sand f. m. c. gr.pb. co. 0 m 5 10 15 20 sandy mud sand gravel lamination structureless sediment climbing ripples trough cross-bedding slump structures ball-and-pillow convolute bedding anastomosing joints thrust fault rubjerg knude formation lønstrup klint formation clayey mud current ripple cross-lamination fig. 27. geological log of sediments and thrust faults in the mr03 thrust sheet in the moserende section (point 1750 in plate 1). 47 garded as a redeposited interglacial fauna by jessen (1931). among the shells are astarte sp., cardium sp., arctica islandica, leda pernula, mya truncata, hiatella arctica; a full list of this diverse fauna is given in jessen (1931, p. 63). an astarte sp. shell (aar-4066) was 14c dated to 43 000 ± 1300 years b.p. (table 2), which must be regarded as close to an infinite age, thus supporting the suggestion of jessen (1931) that these shells represent redeposited interglacial faunas. boundaries. the lower boundary of the formation is placed at the erosional l/r-unconformity capping the lønstrup klint formation. this has a relief of 0.5–1 m and is commonly overlain by an up to 0.5 m thick clast-supported residual gravel bed. in the distal southern part of the rubjerg knude glaciotectonic complex, the l/r-unconformity is located close to sea level, and is a convenient structural reference level. it represents the top level of pre-tectonic sedimentation, and is hence also a reference surface for the construction of the balanced cross-section. the upper boundary is placed at the glaciotectonic unconformity below the kattegat till formation. thickness. the thickness of the formation is about 25 m, but it varies considerably according to local depositional and erosional development. fig. 28. the glaciofluvial ulstrup beds deposited above the l/r-unconformity (l/r-u) on top of the lønstrup klint formation. note the boulder in the lowermost part of the glaciofluvial ulstrup beds indicating the high-energy (upper flow regime) of the meltwater streams that deposited the beds (compare with fig. 20). the divisions on the measuring pole are 20 cm. photograph: may 1998. ulstrup section, 5950 m in cross-section (see plate 1). fig. 29. the glaciofluvial ulstrup beds with ‘fossil frozen’ sand clasts that indicate ground-frozen conditions in the source area of the sand clasts; they were probably derived from the lower part of the rubjerg knude formation farther north. photograph: may 1998. 48 distribution. the rubjerg knude formation was mainly deposited and preserved between the thrust sheets of the rubjerg knude glaciotectonic complex. the formation extends towards the south to the area around nørre lyngby where it was mapped as ‘morænesand’ (moraine sand – sandy till) by jessen (1918, 1931), and it has not been identified south of løkken. the formation is not recognised in the area north of lønstrup, which was mainly covered by ice during deposition of the formation. to the east it can be traced in wells about 10 km inland, where it pinches out due to erosion during the transgression of the younger yoldia sea. the formation probably does not extend out into the north sea to the west since it is largely situated above sea level. age. the formation has an age range of 30 000 – 20 000 years b.p. based on 14c dating of mosses, separated from the organic debris draping the ripple lamination, and twigs and amber layers (houmark-nielsen et al. 1996; table 2). the mosses investigated were transported from a carbonate-rich source area, probably the cretaceous chalk outcrops near limfjorden (fig. 13). the time span for redepositing plant debris is not regarded to exceed hundreds of years, and the age of the formation was thus interpreted to be closer at 29 000 than 30 000 years b.p. (fig. 13; houmark-nielsen et al. 1996). depositional environment. the rubjerg knude formation is interpreted to have been deposited on an outwash plain, which was dissected into smaller piggyback basins during glaciotectonic thrust faulting. during the development of the piggyback basins, deposition was controlled by the propagation of the thrust sheets. north of lønstrup, a large depression is regarded as the hole in a hill-and-hole pair from where the piggyback basins contemporaneous with deposition of the rubjerg knude formation were dislocated to the south during the glaciotectonic deformation. kattegat till formation history. the rubjerg knude formation is truncated by a glaciotectonic unconformity and overlain by the kattegat till formation (fig. 14). the formation was erected by houmark-nielsen (1987) in the areas surrounding the southern part of the kattegat and is interpreted to have been deposited during the weichselian ice advance from norway. subsequent studies have demonstrated that the formation can be identified over much of the northern part of the danish basin (fig. 12; houmark-nielsen 1999, 2003). name. the formation is named after the kattegat strait (fig. 12). type section. the type section is at hundested klint (fig. 12; houmark-nielsen 1987). reference section. two reference sections are defined in the lønstrup klint coastal section, namely the top of the sandrende locality at point 3700 m in plate 1 (figs 19, 30) and the cliff exposure c. 400 m north of the mårup church, at point 500 m in plate 1 (fig. 31). lithology. at the type section, the formation is a grey, clayey till only a few metres thick; the erratic clasts are dominantly crystalline rocks of fennoscandian provenance and palaeozoic limestone. foraminifers and shell fragments in the matrix have been identified as having been derived from the skærumhede group (houmark-nielsen 1987). in the rubjerg knude area, the till is light beige-brown weathering, dark grey and sandy with fineto medium-grained sand in the matrix. erratic pebbles and cobbles occur scattered in the matrix, and indicator pebbles of permian porphyry from the oslo region are common (1–5% of the erratics). in the main part of rubjerg knude cliff section, the formation drapes the glaciotectonic complex; over large areas, it has been subjected to aeolian erosion that has removed the fine-grained matrix and left the erratics as a cobble pavement. north of the mårup church, the formation comprises a shear till with erratics interlayered in a glaciotectonic breccia dominated by shear-deformed clayey mud derived from the top of the skærumhede group (fig. 31). in the northern part of the lønstrup klint cliff section (the ribjerg section), the kattegat till formation is absent. the glacial advance, represented elsewhere by the kattegat till formation, is here recorded only by a glaciotectonic unconformity and an underlying glacitectonite characterised by a dense anastomosing framework of joints penetrating the skærumhede group (fig. 32). boundaries. the lower boundary of the formation is the glaciotectonic unconformity formed by the shear at the base of the advancing norwegian ice. below the unconformity, a glacitectonite 1–2 m thick developed due to shear deformation of the clay and sand in the lønstrup klint and rubjerg knude formations. 49 the upper boundary is the subaerial erosional surface above the 1.5 m thick sandy till, commonly reduced to a 0.25 m thick residual pavement. thickness. the formation is up to 1.5 m thick at rubjerg knude. distribution. the formation has been recognised from the central and northern part of the west coast of jylland and vendsyssel over djursland and sjælland to hven and glumslöv in the western part of skåne, sweden (fig. 12; houmark-nielsen 2003). age. the age of the kattegat till formation is bracketed by the lønstrup klint formation beneath (29 000 years b.p.) and the ribjerg formation above (26 000 years b.p.) (fig. 14, tables 2, 3). the age is estimated to be 27 500 ± 1000 years b.p. (houmark-nielsen 2003). depositional environment. the kattegat till formation is interpreted as a lodgement till. in the area between lønstrup and mårup church, the diamict lithology of the upper skærumhede group suggests that deformation of the substratum below the glaciotectonic unconformity was initiated by mud-mobilisation of water-saturated clay, silt and fine-grained sand. material, including erratic clasts from the lodgement bed along the sole of the ice, dropped into the mudmobilised unit. during the advance of the ice, the mudmobilised zone became consolidated, and sub-horizontal anastomosing joints formed in the substratum. the depositional environment therefore changed from a wet-based glacial advance to an advance over dehydrated or even frozen substratum during the deposition of the kattegat till formation at rubjerg knude. fig. 30. the c. 1 m thick sandy till on top of the sandrende section is referred to the kattegat till formation. the maximum size of the erratic clasts is 25 cm. the marked planar erosion sur face above the till was initially formed by glacial truncation, which subsequently was exposed to aeolian erosion and finally covered by dunes. photograph: july 1993. 50 ribjerg formation new formation history. the ribjerg formation is a new formation proposed for the c. 25 m thick glaciofluvial sand unit that crops out between the northern part of the lønstrup klint clif f section and the northern part of the town of lønstrup. the unit was indicated in the northernmost c. 2 km of the cross-section of jessen (1931), but it was only regarded as part of the main diluvialsand (diluvial sand). since the formation was deposited in the late weichselian between ice advances from norway and central sweden, it might in a glaciodynamic context be correlated with the outwash deposits of the tebbestrup formation in djursland (larsen et al. 1977; pedersen & petersen 1997). name. the formation is named after the hill of ribjerg at lønstrup. type section. the type section is located at the cliff below the ribjerg hill, south-west of lønstrup (figs 13, 33). lithology. the ribjerg formation is characterised by fineto medium-grained sand showing large-scale trough and channel cross-stratification (fig. 33). at fig. 31. a pocket of sandy till overlying a glacitectonite and associated features of subglacial deformation. the till is referred to the kattegat till formation. the locality is situated c. 450 m north of mårup kirke. photograph: july 1994. fig. 32. planar-parallel and elongated anastomosing shear joints that are typical of the glacitectonite at the top of the skærumhede group below the blåunconformity in the ribjerg section. photograph: july 1994. 51 the base of the formation, tabular grey mud clasts (1 × 5–10 cm in cross-section) form a lag deposit in the fine-grained sand. lamination outlined by heavy minerals occurs in the lowermost metre of the formation and current-ripple lamination with mud chips is also present (fig. 33). the lower part of the formation is characterised by horizontal planar laminated finegrained sand interlayered with thin beds (0.1–0.5 cm thick) of clay-draped current ripples. this facies is overlain by 0.5 m thick beds of fineto coarse-grained trough cross-bedded sand. troughs or channels, 2–5 m deep and 10–15 m wide, occur in the middle and upper part of the formation (fig. 111). in the central part of the troughs, the fill shows large-scale crossstratification. towards the margin of the troughs, the beds decrease in thickness and display small-scale mud sand f. m. c. gr.pb. co. 25 20 15 10 5 0 m ribjerg formation peat and dune sand mid danish till formation glaciotectonic-unconformity blå-unconformity skærumhede group x r-990222 x r-990223 x r-990224 sandy mud sand dropstones in mud gravel basal till shells lamination structureless mud clasts climbing ripples trough cross-bedding slump structures ball-and-pillow / convolute bedding water-escape pipes and sand-filled cracks tectonite clayey mud current ripple cross-lamination x aar 4067 fig. 33. sedimentological log of the ribjerg formation (type section) in the northernmost part of the ribjerg section. the formation was deposited above the blå-unconformity on top of the skærumhede group, and it is overlain by the sandy mid danish till formation. the ribjerg formation represents glaciofluvial deposition related to a channel-eroded foreland of an advancing glacier. note the sand-dykes that are interpreted to have formed by discharge of pore water due to high stream velocity. samples collected for optically stimulated luminescence dating are indicated, together with their laboratory numbers (see table 3). 52 current lamination with ripples draped by organic debris. the steeply inclined slopes (up to 30°) of the trough margins strike 88–94°, indicating an east–west current direction. erosional surfaces with slumped beds and pockets of gravel recur every 1 to 3 m (fig. 33). one of the most characteristic features of the formation is the large number of sand dykes and waterescape pillars, which are 5–15 cm wide and can be traced vertically for more than 1.5 m (fig. 34). the uppermost 3 m of the formation comprises thick gravel beds just below the flow till related to the mid danish till formation. boundaries. the lower boundary of the formation is defined at the erosional unconformity forming the top of the kattegat till formation or, where the erosion level penetrates deeper, the lønstrup klint formation. in the northern part of the lønstrup klint section, the unconformity on top of the ‘lille blå’ forms the lower boundary. the upper boundary is placed at the base of the flow till that forms the lower part of the mid danish till formation (figs 14, 33). thickness. the formation is about 25 m thick. distribution. the formation is only recognised in the fig. 34. sand dyke intruded in the glaciofluvial succession of the ribjerg formation. the sand dykes are interpreted to have formed by pore-water discharge from the sediment due to high velocity current flux through the channels. photograph: july 1994. ribjerg formation, type section. 53 vicinity of lønstrup and towards vennebjerg to the east. it is inferred to have been deposited over a larger area of north-western vendsyssel, which is now covered by the vendsyssel formation (see below). age. three samples were collected from the lower, middle and upper part of the formation for optically stimulated luminescence dating (r-990222, r-990223, r-990224; table 3); these samples indicate an age of 26 000 – 25 000 years b.p. depositional environment. the formation was deposited in fluvial channels cut by westward-flowing meltwater. the sand dykes and water-escape pillars are indicative of high pore-water pressure due to rapid deposition and very fast meltwater flux through the channels. the outwash deposits are interpreted as a valley sandur that formed in the depression resulting from the hole left in the hinterland of the rubjerg knude glaciotectonic complex. the source of the meltwater was the ice margin of the advancing ice from central sweden in the late weichselian. mid danish till formation history. the ribjerg formation of the lønstrup klint section is overlain by a 3 m thick grey-brown till that is referred to the mid danish till formation (houmark-nielsen 1987, 1999, 2003). the formation was erected by houmark-nielsen (1987) to encompass tills deposited in the southern and central part of denmark during the weichselian ice advance from central sweden. it is known from the main part of the danish basin east and north of the main stationary line (fig. 12), and the records of its distribution in the northern part of denmark have recently been summarised by houmark-nielsen (1999, 2003) (figs 1, 12). name. the name of the formation reflects the prominent nature of this surface deposit in central (mid) denmark (fig. 12). type section. the type section is at ristinge klint on the island of langeland (houmark-nielsen 1987). reference section. a reference section is herein defined at ribjerg, sw of lønstrup (figs 13, 33). lithology. at the type locality, the mid danish till formation is a 5–8 m thick unit with at least two boulder pavements displaying ne–sw orientated glacial striation (sjørring et al. 1982). the formation is a grey to brown mostly clayey massive till with about 50% crystalline erratics of fennoscandian provenance; indicator clasts from the central eastern part of sweden (kinne-diabase) and from jurassic sedimentary rocks situated offshore in the kattegat (pedersen & petersen 1997) are abundant (houmark-nielsen 1987). in the reference section at ribjerg, the formation is table 3. optically stimulated luminescence dates on quartz, rubjerg knude and nørre lyngby, vendsyssel, northern denmark stratigraphic unit locality lab. id no. material age ka b.p. dose+ (gy) w.c.‡ (%) ref.* vendsyssel fm nørre lyngby r-829202a marine clay 16 ± 1 38.9 ± 1.3 30 (1) vendsyssel fm nørre lyngby r-829203 marine clay 17 ± 2 46.0 ± 0.8 29 (1) ribjerg fm ribjerg r-990224 fluvial sand 25 ± 2 45.3 ± 0.8 25 (3) ribjerg fm ribjerg r-990223 fluvial sand 26 ± 1 52.6 ± 1.5 26 (3) ribjerg fm ribjerg r-990222 fluvial sand 26 ± 1 53.1 ± 1.2 20 (3) lønstrup klint fm sandrende r-970204 fluvial sand 29 ± 2 57.6 ± 1.8 21 (2) stortorn fm ribjerg r-970203 lacust. sand 30 ± 2 65.0 ± 1.8 25 (2) + equivalent gamma dose. ‡ water content (saturation). * references: 1: strickertson & murray (1999); 2: houmark-nielsen (2003); 3: this study. 54 a yellow-brown weathering, grey-brown, sandy till comprising a lower stratified unit and an upper massive unit (fig. 35). the lower unit comprises laminated to finely bedded, matrix-supported diamictite with scattered pebbles. the matrix is fine-grained sand and the lamination and bedding are slump folded with n– s-trending fold axes and e-dipping axial planes, indicating a westward flow direction. the upper unit is a massive, structureless matrix-supported diamictite. erratic pebbles and cobbles are abundant and the till has a pronounced a-axis clast fabric dipping at low angles (c. 3°) towards the east (100°), indicating a shear transport direction towards the west. boundaries. in the reference section, the lower boundary of the formation is placed at the depositional conformity on top of the ribjerg formation where planar horizontal gravel beds are overlain by slumpfolded diamictites dominated by debris flow layers. the upper boundary is an erosional unconformity separating the diamictites from silt-streaked muds at the base of the vendsyssel formation. thickness. the formation reaches a thickness of more than 10 m at the type section at ristinge klint, but it is only 3 m thick at ribjerg in the reference section. distribution. in the rubjerg knude area south of rifig. 35. the sandy till that overlies the ribjerg formation is divided into a lower flow till and an upper lodgement till. the flow till is characterised by slump-folded debris flow lamination indicating flow from east to west. the till unit is referred to the mid danish till formation, which was deposited by the ice advance from the east, probably about 24 000 b.p. photograph: may 1985; notebook for scale. ribjerg formation, type section. 55 bjerg, where the ribjerg formation is absent, it has not been possible to differentiate occurrences of the mid danish till formation from the older kattegat till formation. however, the distribution is well documented throughout the danish basin east and north of the main stationary line (figs 1, 12; houmark-nielsen 1999, 2003). age. the age of the mid danish till formation is bracketed by the age of the ribjerg formation beneath (26 000 years b.p.) and the vendsyssel formation above (16 000 years b.p.) (fig. 14, tables 2, 3). the age is estimated to be 24 000 – 20 000 years b.p. (houmark-nielsen 2003). depositional environment. the lower unit is interpreted as a flow till deposited as debris flows from an ice margin to the east, prior to the ice advance towards the west. the upper unit is interpreted as a lodgement till deposited at the sole of the ice during the ice advance from central sweden towards the main stationary line situated in the central part of the north sea (fig.12). vendsyssel formation new formation history. north and south of rubjerg knude, the mid danish till formation is overlain by a succession of glaciomarine heteroliths, which are here defined as the vendsyssel formation. these deposits were mapped by jessen (1899), who related them to deposition in the younger yoldia sea in vendsyssel. jessen (1918) regarded the various facies of the vendsyssel formation as four stratigraphic units named the lower saxicava sand, the yoldia clay (usually prefaced ‘youngaeolian sand vendsyssel formation lønstrup klint formation glacitectonite 0 m 5 10 15 sandsilt m.f.si. c.clay gr. co. tectonite climbing ripples current ripple cross-lamination sand wave ripple lamination lamination trough cross-bedding gravel sandy mud clayey mud storm sand bed bioturbation shells imbricated mud clasts ball-and-pillow convolute bedding fig. 36. sedimentological log of the vendsyssel formation at the type section at stensnæs. the section is located at point 5120 in the cross-section (plate 1); the base of the log is 15 m a.s.l. 56 er’ to distinguish it from the older yoldia clay), the upper saxicava sand and the zirphaea beds. name. the formation is named after the region of vendsyssel in north denmark (figs 12, 13). type section. the type section is the coastal cliff section at stensnæs c. 1 km north of nørre lyngby in the central part of the west coast in vendsyssel (figs 13, 36). reference sections. two coastal cliff sections, north and south of rubjerg knude, are defined as reference sections. the locality to the north is the coastal clif f c. 500 m south of lønstrup, where heteroliths characterised by hiatella burrows crop out (fig. 37). to the south, the coastal cliff at nørre lyngby (north and south of the ramp leading down to the beach) probably gives the thickest accessible outcrop of the formation (fig. 38). this locality is furthermore close to the reference well dgu no. 8.137, where the maximum thickness of the formation is recorded (lykkeandersen 1987). lithology. two main lithologies dominate the formation: dark bluish-grey, clayey mud in the lower part and yellowish weathering light grey stratified heteroliths in the upper part. at the base of the formation, coarse-grained sands and gravels overlie the erosional unconformity above the mid danish till formation or older deposits (figs 36–38). the unit referred to as the lower saxicava sand by jessen (1918) is less than 2 m thick and is only present locally. accumulations of shell debris occur in places. in general, marine clayey mud forms the lower c. 6 m of the formation resting tectonite basal till aeolian sand vendsyssel formation kattegat till formation lønstrup klint formation glacitectonite 0 m 5 10 15 sandy mud sand gravel lamination wave ripple lamination storm sand bed bioturbation shells trough cross-bedding slump structure clayey mud sand m.f.si. c.clay gr.pb. co. fig. 37. sedimentological log of the vendsyssel formation at the reference section, situated halfway between lønstrup and mårup church at point 500 in the cross-section (plate 1); base of the log is 10 m a.s.l. the slump structure recognised in the lower part of the section (at about 4 m) is interpreted to have been produced by a grounding iceberg. the abundant shells in the section are hiatella arctica and the bioturbation was due to the infaunal activity of these molluscs (see fig. 41). 57 directly on the lower erosional boundary (figs 36– 38); dropstones, locally to boulder size, occur in the lower part of the clayey mud unit (figs 36, 37). the unit is highly impermeable such that groundwater wells out at the top of the clayey mud outcrops, often obscuring the exposures of the basal lithologies. above the clayey mud unit, horizontal stratified heteroliths form a unit 6–12 m thick. in places, the heteroliths grade into sandy mud characterised by wave ripple lamination (fig. 40). dark grey laminated mud is interbedded with fine-grained sand beds up to 10 cm thick in which wave ripple lamination is common. at the reference section, south of lønstrup, the heteroliths are intensively bioturbated by vertical trace fossils produced by hiatella arctica and the shells are often preserved in life position (fig. 41). the reference section at nørre lyngby is located in a half-graben structure with the steepest normal fault (dipping c. 60°s) situated north of the village (lykkeandersen 1992). south of nørre lyngby, the erosional unconformity below the vendsyssel formation dips 5–8° to the north. the beds above the unconformity are characterised by sedimentary breccias of mud clasts probably derived from the lønstrup klint formation (fig. 38). in the middle part of the formation, the beds are displaced by synsedimentary faulting (fig. 38) indicating that the half-graben formed during the deposition of the vendsyssel formation. aeolian sand vendsyssel formation kattegat till formation lønstrup klint formation sandsilt m.f.si. c.clay gr. co. 0 m 5 10 15 20 dropstones in mud tectonite climbing ripples current ripple cross-lamination basal till sandy mud sand lamination trough cross-bedding clayey mud imbricated mud clasts ball-and-pillow / convolute bedding normal fault thrust fault fig. 38. sedimentological log of the vendsyssel formation at the reference section, situated c. 350 m south of nørre lyngby. the beds rich in imbricated mud clasts reflect the tectonically active nature of the half-graben in which the section is located; the synsedimentary tectonic activity is further documented by the intraformational normal and thrust faults that occur in the lower half of the formation at nørre lyngby. 58 in the area north-east of hirtshals, the clayey mud is overlain by coastal sands, the so-called zirphaea beds (jessen 1918), and in the eastern part of vendsyssel the uppermost part of the formation comprises coarsegrained sand and gravel deposited in a large spit system (nielsen et al. 1988). fossils. the fossils characteristic of the formation are the molluscs portlandia arctica, hiatella arctica and zirphaea crispata. boundaries. the lower boundary is the erosional unconformity on top of the mid danish till formation or older deposits. the upper boundary is the top surface of the landscape upon which locally lie terrestrial deposits such as the allerød peat beds in the nørre lyngby bog (jessen & nordmann 1915), the boreal peat at martørv bakker (jessen 1931) and recent aeolian sands (fig. 14). thickness. the formation is c. 16 m thick at the outcrops along the coastal cliff. the formation may reach fig. 39. mud-dominated heteroliths in the lower vendsyssel formation. the light-coloured silt to very fine-grained sand beds show grading and wave ripple cross-lamination. this unit was formerly referred to as the younger yoldia clay (jessen 1918, 1931). photograph: september 2004. fig. 40. sand-rich heteroliths in the upper vendsyssel formation showing wave ripple cross-lamination. this sandrich heterolithic unit was formerly referred to as the saxicava sand (jessen 1918, 1931). tape divisions in centimetres. photograph: september 2004. 59 a thickness of up to about 25 m in the central part of vendsyssel (see fig. 125). distribution. the flat agricultural land in the vendsyssel area, lying 10–40 m above sea level, defines the top of the vendsyssel formation, and thus can be regarded to represent the fossil seabed of the younger yoldia sea. depositional environment. the formation reflects the establishment of marine conditions in the vendsyssel area after the melting back of the scandinavian ice cap in the kattegat–vendsyssel–skagerrak region. the palaeoenvironmental development may be described in terms of six events (richard 1996): the first event is represented by the erosional unconformity formed immediately after deglaciation. the second event was a rapid transgression with the establishment of a c. 60 m deep arctic marine environment. in the third event, a high sea-level stand prevailed during deposition of the clayey mud. events four to six are stages of forced regression due to the isostatic uplift in the area, but with fluctuations due to eustatic sea-level rise. age. the age of the vendsyssel formation ranges from 17 000 to 14 500 b.p. (tables 2, 3; tauber 1966; krog & tauber 1974; knudsen 1978; abrahamsen & readman 1980; aaris-sørensen & petersen 1984; nielsen et al. 1988; richard 1996; houmark-nielsen 2003). fig. 41. at the cliff section north of mårup church, the sandy mud is often highly bioturbated; in places, the shells of the bivalve hiatella arctica are found in life position in the escape trace fossils. photograph from the middle part of the section in fig. 37. photograph: september 2004. colgan et al. 2025: geus bulletin 59. 8387. https://doi.org/10.34194/ft7sk529 1 of 17 research article traverse route from pituffik to warming land, north greenland william colgan1* , dominik fahrner1 , gregor luetzenburg1 , anja løkkegaard1 , eva bendix nielsen1 , penelope how2 , anja rutishauser1 , anne solgaard1 , robert s. fausto1 , kristian k. kjeldsen1 , christopher shields1, henrik spanggård1, patrick smith3, joseph a. macgregor4 , andreas p. ahlstrøm1 nanna b. karlsson1 1department of glaciology and climate, geological survey of denmark and greenland (geus), copenhagen, denmark; 2department of glaciology and climate, geological survey of denmark and greenland (geus), nuuk, greenland; 3polar field services, littleton, colorado, usa; 4cryospheric sciences laboratory, nasa goddard space flight center, greenbelt, usa abstract we examine the feasibility of an overland motorised traverse from pituffik to greenland’s oldest ice outcrop in warming land, north greenland. we assess a 778 km overland traverse that departs pituffik via the nunatarssuaq take-off ramp, which is an alternative to the more frequently used, but more heavily crevassed, thule take-off ramp. the traverse route includes brief sea ice and primitive road conditions, each c. 4% of the route length, and a lengthy ice sheet segment (c. 92% of the route length). this study outlines challenges for each of these traverse segments, including primitive road conditions and snow cover, seasonality of extreme cold conditions (air temperatures below –30°c), seasonality of surface melting and softening (air temperatures above 0°c), sea-ice thickness and potential ridging hazards and ice dynamics and potential crevasse hazards. ongoing work is required for annual vetting of the traverse route to ensure operational safety. the optimal operational window for such a traverse is departing pituffik in mid-april and returning in mid-may. in comparison to aircraft-supported ice-sheet fieldwork, scientific traverses offer the opportunity for more intensive ground-based science, while significantly reducing carbon emissions. based on previously reported traverse fuel consumptions, a ground traverse from pituffik to warming land would use 90% less fuel than aircraft-supported fieldwork. this assessment underscores the potential for sustainable ground-based access to greenland’s oldest ice outcrop and other science sites within the region. *correspondence: wic@geus.dk received: 12 dec 2024 revised: 23 may 2025 accepted: 17 jun 2025 published: 20 oct 2025 keywords: traverse, ice sheet, pituffik, warming land abbreviations cen: camp century weather station gc-net: greenland climate network geus: geological survey of denmark and greenland hum: humboldt weather station ngrip: north greenland ice core project nsf: national science foundation nuto: nunatarssuaq take-off pdd: positive degree days promice: programme for the monitoring of the greenland ice sheet thu_l: thule lower tuto: thule take-off usace: u.s. army corps of engineers geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: harold lovell (university of portsmouth, uk) reviewed by: zoe courville (cold regions research and engineering laboratory, usa), lenneke jong (australian antarctic division, australia) and one anonymous reviewer funding: see page 15 competing interests: see page 15 additional files: see page 16 greenland’s oldest ice outcrop ice cores, drilled vertically into earth’s ice sheets and glaciers, have provided valuable information on past climates (jouzel & masson-delmotte 2010). retrieving precious ancient ice, however, comes at great cost and typically takes several years. collecting so-called ‘horizontal ice cores’ seeks to make use of the fact that ancient ice can also resurface at land-terminating margins. while this approach has long offered the tantalising possibility to recover nearly unlimited volumes of paleo ice of a desired age, reliably assessing surface ice age has been a tremendous practical challenge to such efforts (reeh et al. 2002). the few attempts to collect horizontal chronologies of ice, however, have benefited from more detailed analyses than are possible with volume-limited conventional ice core samples recovered from depth (petrenko et al. 2006; kurbatov et al. 2010). the age and extent of last glacial period ice outcropping from the icesheet ablation area has recently been mapped in north greenland using visible spectrum satellite imagery linked to ice-core stratigraphy and chronology (macgregor et al. 2020). alternating light and dark bands within this last glacial period ice outcrop correlate with periods of higher and lower dust https://doi.org/10.34194/ft7sk529 https://orcid.org/0000-0001-6334-1660 https://orcid.org/0000-0002-7895-1557 https://orcid.org/0000-0001-5443-7572 https://orcid.org/0000-0002-1947-5773 https://orcid.org/0000-0002-6888-6675 https://orcid.org/0000-0002-8088-8497 https://orcid.org/0000-0002-1819-8014 https://orcid.org/0000-0002-8693-620x https://orcid.org/0000-0003-1317-8185 https://orcid.org/0000-0002-8557-5131 https://orcid.org/0000-0002-5517-2235 https://orcid.org/0000-0001-8235-8070 https://orcid.org/0000-0003-0423-8705 mailto:wic@geus.dk https://creativecommons.org/licenses/by/4.0/deed.ast colgan et al. 2025: geus bulletin 59. 8387. https://doi.org/10.34194/ft7sk529 2 of 17 geusbulletin.org concentration in the north greenland ice core project (ngrip) ice core. this reflects greenland stadials and interstadials associated with dansgaard–oeschger events. in warming land, north greenland, the alternating light and dark ice bands allow the ablation area ice to be potentially mapped to at least 48 000 years before the present time (macgregor et al. 2020). this makes the warming land ice-sheet ablation area a high priority target for paleoclimatic sampling. we describe general science plans for the warming land ablation area towards the end of this report. the warming land ablation area containing greenland’s oldest ice outcrop is located between sermeq konrad steffen and sermeq niels reeh. this is an exceptionally logistically challenging location to reach. while knud rasmussen’s second thule expedition (1917–1918) and svend lauge koch’s bicentenary jubilee expedition (1920–1923) both crossed the ice-sheet ablation area in warming land, only one modern scientific party, a 1984 geus (geological survey of denmark and greenland) team led by niels henriksen, has undertaken glaciological investigations there (henriksen et al. 2009; dawes 2022). located c. 530 km from the airport at qaanaaq to the south-west and c. 580 km from the airport at station nord to the east, the site is difficult to reach by aircraft due to the long distances over which aircrafts must carry their own return fuel. aircraft logistics are therefore poorly suited for sustaining scientific investigations at greenland’s oldest ice crop. motorised ice-sheet traverses an alternative to using aircraft is using ground vehicles for overland traverses on the ice sheet (fig. 1). between 1959 and 1967, traverses of the u.s. army corps of engineers (usace) regularly departed the pituffik space base (formerly thule air base) to support camp century inland on the ice sheet. these traverses primarily ascended the ice sheet at thule take-off (tuto) ramp, south of pituffik (davis 1967). between 2008 and 2016, tuto ramp was utilised by the u.s. national science foundation (nsf) to operate 2200 km round-trip resupply traverses between pituffik and summit (lever & weale 2011). recent increases in crevasse extent c. 40–45 km inland from tuto ramp have effectively ended the operational use of this route. this inland crevasse area is located where the traverse route follows along a relatively narrow local flow divide between ullip sermia (harald moltke bræ) flowing into uummannap kangerlua (wolstenholme fjord) to the north and a family of tidewater glaciers draining into puisillip kangerlua (de dødes fjord) to the south. before selecting tuto ramp as the primary ice-sheet ascent to camp century, the usace also trialled ascending the ice sheet at nunatarssuaq take-off (nuto) ramp, on the north side of uummannap kangerlua (wolstenholme fjord; fig. 2). this land–ice transition had been previously used by the british north greenland expedition (hamilton et  al. 1956). while nuto ramp was located at an elevation similar to tuto ramp, the steeper ice-sheet surface there presented a shorter ablation area to traverse. the tuto and nuto icesheet ascents, after differing lengths of over ice travel, converged at ~1200 m elevation on the ice sheet, at a site named sierra. the seasonal constraint of crossing the sea ice from pituffik to nuto ramp ultimately compelled the usace to construct a gravel road across the ablation area at tuto ramp to permit year-round icesheet access. nuto ramp has fallen almost entirely out of use since that time. during the 1950s and 1960s, the usace also performed extensive testing of the trafficability of different vehicle types on the ice sheet in north-west greenland (nuttall & finelli 1955; rula 1960; mellor 1963; schreiner 1965). this testing provides a wealth of information that continues to be relevant today. in these investigations, the usace clearly identified that vehicle performance is dependent on snow type and wetness, with all vehicles performing best on dry fine-grained snow. while wheeled vehicles perform adequately in the ice sheet’s interior accumulation area, tracked vehicles consistently outperformed wheeled vehicles, especially when the snow is either deep and soft or wet and saturated. the usace reports of the era therefore make the specific recommendation that tracked vehicles with ground pressure under 6 psi (41 kpa) provide the most effective means of ice-sheet transportation in north-west greenland across all conditions. pittufik to warming land here, we look at the technical and logistical feasibility of supporting scientific investigations at greenland’s oldest ice outcrop, in warming land, using an overland traverse from pituffik. as an alternative to the heavily crevassed tuto route, we assess a route that ascends the ice sheet at nuto ramp (fig. 2). we provide an overview of the route, and describe specific challenges of the sea ice, the primitive road and the ice sheet segments of the traverse (fig. 3). we also describe general science plans for working in the warming land ice-sheet ablation area. lastly, we assess the differences in carbon emissions associated with accessing warming land via aircraft and traverses. the traverse route from pituffik to warming land, via nuto ramp, is 778 km long one way. it begins at the north star bugt sea-ice access ramp at pituffik space base, located at sea level at the harbour. the first 29 km of the route crosses sea ice around uummannaq (dundas fjeld) into uummannap kangerlua (wolstenholme fjord), https://doi.org/10.34194/ft7sk529 https://geusbulletin.org/ colgan et al. 2025: geus bulletin 59. 8387. https://doi.org/10.34194/ft7sk529 3 of 17 geusbulletin.org fig. 1 overview placing the traverse route assessed in this study in the context of past traverses on the greenland ice sheet. past traverses: aiegi: american-icelandic expedition on the greenlandic icecap (stokholm et  al. 2021); greentracs: greenland traverse for accumulation and climate studies (meehan et al. 2021); firncover: greenland firn compaction verification and reconnaissance (macferrin et al. 2022); grlt: greenland ice sheet traverse (hawley et al. 2014); ucph: university of copenhagen (karlsson et al. 2020); ngt: north greenland traverse; parca: program for arctic regional climate assessment (thomas et al. 1998); egig: expédition glaciologique internationale au groenland (hofmann 1964); usace: u.s. army corps of engineers (usatb 1960; benson 1962); bnge: british north greenland expedition (hamilton et al. 1956). this compilation of previous motorised traverses is likely not comprehensive. stars indicate place names referred to in the text. hum: humboldt. cen: camp century. nem: north greenland eemian ice core (neem). sigma-a: snow impurity and glacial microbe in the arctic site. dye-2: disused radar station. aiegi (2018) greentracs (2016–2017) firncover (2015–2017) grit (2008–2016) ucph (2007; 2015) ngt (1993–1995) parca (1994) egig (1951–1960; 1965–1967) usace (1952–1955; 1960) bnge (1952–1954) this study (primary) this study (secondary) ice mask ice-free land ice covered https://doi.org/10.34194/ft7sk529 https://geusbulletin.org/ colgan et al. 2025: geus bulletin 59. 8387. https://doi.org/10.34194/ft7sk529 4 of 17 geusbulletin.org fig. 2 overview of the study area. a: highlighting the tuto and nuto ramps, as well as place names in the vicinity of pituffik (klimadatastyrelsen 2024a). the black dashed line represents the traverse routes. b: mile markers on the overland traverses from the nuto and tuto ice-sheet approaches (benson 1962). reproduced from the u.s. corps of army engineers. fig. 3 overview of the traverse route from pituffik to warming land, with ice-sheet surface velocities overlaid on a visible spot 6 satellite image mosaic of 10 m spatial resolution for the region from july 2020 (solgaard & kusk 2022; solgaard et al. 2021; klimadatastyrelsen 2024). potential route extensions linking secondary ice-sheet science sites in the area are also shown. secondary ice-sheet science sites are denoted thus: hum for humboldt, cen for the camp century, nem for the north greenland eemian ice core (neem) and sigma-a for the snow impurity and glacial microbe in the arctic site (matoba et al. 2018; vandecrux et al. 2023). https://doi.org/10.34194/ft7sk529 https://geusbulletin.org/ colgan et al. 2025: geus bulletin 59. 8387. https://doi.org/10.34194/ft7sk529 5 of 17 geusbulletin.org making landfall again at nunatarsuup nuua (kristiansen point). the subsequent 32 km of the route (km 29 to 61) then traverses a primitive road constructed by the usace in 1954. this primitive road rises from sea level to the margin of the ice sheet at 552 m elevation. this primitive road reflects c. 4% of the total traverse length, but >25% of the total traverse elevation climb. the remaining 717 km of the route traverses the ice sheet, terminating in the warming land ablation area. on the ice sheet, the proposed route is selected with the primary goals of minimising altitudinal variations and ice velocities (fig. 4). the route ascends the ice sheet via the centreline of nuto ramp and continues climbing a local flow divide to where it intersects the tuto ramp approach route at km 119 (1443 m elevation). to minimise crossing high ice velocities and strain rates associated with tributaries of ullip sermia (harald moltke bræ), our proposed nuto ascent bears north of the usace nuto ascent. consequently, the modern sierra, or intersection of the tuto and nuto approaches, is c. 16 km further east of the usace sierra (mile 1–10 instead of mile 1–0; fig. 2). from this modern sierra crossroads, the route then follows the local ice-flow divide upwards to camp century weather station at km 221 (cen; 1895 m elevation). the route then turns north-east and follows just below the 2000 m contour, passing the humboldt weather station at km 412 (hum; 1950 m elevation), and continues east, past the inland ice-stream onset of petermann gletsjer (fig. 3). it then turns north to descend a local ice-flow divide into the land-terminating ice-sheet ablation area in warming land at km 778 (902 m elevation). a shapefile of this traverse route is available in the online supplementary material associated with this report (supplementary file s1). sea-ice segment the 29 km sea-ice portion of the traverse connects between pituffik and nunatarsuup nuua (kristiansen point). we assess sea-ice characteristics along this portion of the traverse route during the 2014–2024 period using sea ice classification charts from the danish meteorological institute (dmi 2024). we extract sea ice extent from landsat 8 satellite imagery with a 30 m resolution. the dmi sea ice charts are determined by the interpretation of blended synthetic aperture radar imagery, optical imagery, infrared temperatures and passive microwave imagery. together, these ice type and extent data show that sea ice thickness and extent demonstrate a consistent seasonality within uummannap kangerlua (wolstenholme fjord; fig. 5, table 1). land-fast ice is present within the fjord each winter until the open water retreats inside appat (saunders ø) at the mouth of uummannap kangerlua. this breakup of the land-fast ice generally occurs in june (weeks 23–27) but has been observed to occur in may (weeks 18–22) at least once during the 2014–2024 period. the traverse route plans to utilise the maintained land-to-sea transition at pituffik harbour (fig. 6). it is, however, also possible to transition from land to sea at dundas village, north of the pituffik space base, inside uummannap kangerlua. using the dundas transition point, which remained within the may land-fast ice perimeter in all years during fig. 4 feature of the traverse routes. a: ice surface velocity along the tuto and nuto traverse routes to warming land derived from a 5-year average (2016–2021) based on offset tracking of sentinel-1 data (solgaard & kusk 2022; solgaard et al. 2021). b: surface elevation along both routes (porter et al. 2023). vertical lines mark the location of the promice/gc-net weather stations along the routes. these stations are denoted thus: thu_l for thule lower, cen for camp century and hum for humboldt (vandecrux et al. 2023; how et al. 2022). https://doi.org/10.34194/ft7sk529 https://geusbulletin.org/ colgan et al. 2025: geus bulletin 59. 8387. https://doi.org/10.34194/ft7sk529 6 of 17 geusbulletin.org 2014–2024, may mitigate early sea-ice breakup in the vicinity of appat. uummannap kangerlua is rarely filled with land-fast ice throughout june. sea ice of at least 30 cm thickness is present along the traverse route throughout april (fig. 5). this includes the dmi classifications of ‘fast ice’ and assorted ‘winter ice’ types, which itself includes ‘thin’ – (30–70 cm), ‘middle’ – (70–120 cm) and ‘thick’ – (>120 cm) types (dmi 2024). around mid-may (week 19), thinner sea-ice classifications, including ‘young’, ‘grey’, ‘thin’ and ‘new’, become increasingly present, as the fjord regularly transitions to open water by late june. within this range of winter sea-ice thicknesses, canada’s government of the northwest territories recommends wheeled vehicle loads, ranging from 3600 kg at 30 cm thickness to 57 600 kg at 120 cm thickness (gnwt 2015). this assumes a natural ice-pressure factor of 4 kg/ cm2 or 392 pa, which may be exceedingly conservative for the low-ground pressure vehicles suitable for traversing the ice sheet. in any case, the acceptable load for the seaice traverse can be most accurately estimated by verifying the thickness of the sea-ice route each year. this can be done via an ice-penetrating radar survey with accompanying ice coring (masterson 2009; barrette 2015). similar to the ice-sheet segment of the traverse, trafficability of the sea ice segment can be hindered by positive (above-freezing) air temperatures, which warm and soften the snowpack. near sea-level meteorological records observed at pituffik suggest that the average date of thaw onset, or positive air temperatures, is 17 may over the 1995–2024 period (menne et  al. 2023; jensen et  al. 2025). the earliest date of thaw onset during this period has been 18 march (2013), and the latest date of thaw onset during this period has been 15 june (2022). traverse operations on the sea ice should therefore ideally occur prior to mid-may, to avoid softening snowpack, and certainly before the end of may, to avoid spring breakup of the land-fast sea ice. 2014 2015 2017 2016 2018 2019 2020 2021 2022 2023 2024 year 2014 2015 2017 2016 2018 2019 2020 2021 2022 2023 2024 year april (week 14–17) may (week 18–22) june (week 22–27) d ev el op m en t s ta ge fast ice winter ice (30–220 cm) young ice (10–30 cm) thin ice (<10 cm) new ice ice of land origin open water winter ice (30–220 cm) young ice (10–30 cm) week 14 16 18 20 22 24 26 a b c d fig. 5 sea-ice conditions in the vicinity of the traverse routes. a: sea-ice development stage assessed in sea-ice charts of the danish meteorological institute (dmi 2024). b–d: sea-ice extent delineated in landsat 8 visible imagery of 15 m resolution for april, may and june during the period 2014–2024. acquisition dates listed in table 1. yellow line: sea-ice segment of the traverse. table 1 acquisition dates of the available landsat 8 cloud-free imagery used for mapping of sea-ice extent during the 2014–2024 period in fig. 5. year april may june 2014 8, 10, 13, 17 19, 24, 26 2, 6, 13, 25 2015 2, 20, 25, 27, 29 4, 18, 20, 22, 27, 29 28, 30 2016 2, 6 6, 15, 22, 24, 29 5, 6, 10, 27, 29 2017 5, 7, 14 2, 11, 16, 30 5, 17, 23, 29, 30 2018 3, 17 10, 14, 26 4, 6, 11, 15, 27, 28 2019 6, 10, 15, 21, 29, 31 7, 13, 27, 29 2020 6, 24 10, 14, 17, 24 6, 8, 13, 18, 20, 27 2021 9, 25 2, 4, 6, 26 5, 7 2022 3, 21 9, 13, 14, 16 1, 15, 17, 22, 24, 30 2023 1, 6, 8 ,10 19 1, 2, 4, 24, 26, 29 2024 10 5, 19, 21, 26 4, 10, 15, 19, 20, 22, 26, 28, 29 https://doi.org/10.34194/ft7sk529 https://geusbulletin.org/ colgan et al. 2025: geus bulletin 59. 8387. https://doi.org/10.34194/ft7sk529 7 of 17 geusbulletin.org primitive road segment the 32-km primitive road segment traverses the nunatarsuaq plateau, from the ocean shore at nunatarsuup nuua (kristiansen point) to the ice-sheet margin at nuto ramp (fig. 7). fortunately, the land-fast nature of sea ice within uummannap kangerlua (wolstenholme fjord) minimises the development of ice ridges associated with sea-ice thrusting along the shoreline. the sandy spit at nunatarsuup nuua has shallow relief and presents two nearly orthogonal beach aspects. this offers some protection against sea-ice ridging by both minimising local formation and providing perpendicular approaches. from nunatarsuup nuua, the route ascends the primitive road to the ice-sheet margin at the nuto ramp. the primitive road appears to have been actively maintained by the usace between its 1954 initial use and the subsequent 1960 completion of an improved ramp at tuto in support of camp century operations (wes 1959; davis 1967). the nuto approach has received virtually no maintenance since then, however, although recreational vehicles from pituffik are known to have used the route in recent years. to evaluate the current condition of the nuto access road, we analysed a mosaic of spot visible imagery and individual scenes tandem-x radar imagery, with spatial resolutions of 1.6 m and 6 m, respectively. in the first 10 km from nunatarsuup nuua, the road climbs more than 500 m in elevation and then a b fig. 6 views to the traverse. a: oblique aerial photograph of land-fast ice within uummannap kangerlua (wolstenholme fjord) on 22 may 2018. naajat sermiat (chamberlin glacier) in the background. photo: william colgan. b: photograph of the land-to-sea transition maintained at pituffik harbour on 17 may 2022. uummannaq (dundas fjeld) in the background. photo: william colgan. the location of the photos is depicted in fig. 2. the black dashed line represents the traverse route. https://doi.org/10.34194/ft7sk529 https://geusbulletin.org/ colgan et al. 2025: geus bulletin 59. 8387. https://doi.org/10.34194/ft7sk529 8 of 17 geusbulletin.org follows natural topography on the nunatarsuaq plateau. nearly the entire road length appears to be in good condition, with the raised and graded material of the roadbed intact and visible from space (fig. 7). at seven stream crossings, however, the roadbed has clearly been eroded. such washouts of road material are regularly reported during nuto ramp’s operational period (wes 1959). wind sheltering can cause drifting snow to preferentially accumulate in topographic depressions, due to aeolian deposition associated with decreased windspeeds in sheltered depressions (mott et  al. 2018). geus has operated on the tuto access road most springs since c. 2010. in our experience, windblown spring snowpack collects on the road where it forms a local depression relative to surrounding topography. we therefore expect that windblown spring snowpack will preferentially collect in washouts on the nuto access road. spring snow coverage, however, is far from uniform on the thule peninsula, both through time and space (fig. 8). from operational experience at tuto access road, we expect that the majority (c. 75%) of the access road remains effectively snow-free, due to its elevated relief, with only a minority (c. 25%) of the access road having sufficiently depressed local relief to develop a windblown snowpack. if the nuto access road shares these general characteristics, then depressed stream crossings will likely be snow-filled each spring. the majority of the primitive road traverse will also therefore take place on gravel, rather than snow. ice-sheet segment the 717 km long ice-sheet segment begins with the ascent of nuto ramp and ends in the ice-sheet ablation area in warming land (nobles 1960). this longest portion of the traverse route has two main challenges, trafficability of the surface and crevasse safety hazards. the trafficability of the ice-sheet surface is highly dependent on cold and firm snow conditions, as warm and soft snow can significantly hinder vehicle performance. surface melting can also present obstacles such as supraglacial lakes and streams. we assess the likelihood of encountering melting conditions using observations from the promice (programme for the monitoring of fig. 7 map of a primitive road on the nunatarsuaq plateau between nunatarsuup nuua (kristiansen point) and the ice-sheet margin. a: topographic map of the entire road length with suspected washouts of the roadbed indicated in red (klimadatastyrelsen 2024a). b: visible image at landfall, nunatarsuup nuua. c: visible image of the nuto ice ramp. d: detailed visible image of a road segment in good condition. e: detailed visible image of a road segment likely eroded by a river crossing. all visible satellite images are spot 6 images from july 2020 (klimadatastyrelsen 2024b). the black dashed line represents the traverse route. https://doi.org/10.34194/ft7sk529 https://geusbulletin.org/ colgan et al. 2025: geus bulletin 59. 8387. https://doi.org/10.34194/ft7sk529 9 of 17 geusbulletin.org the greenland ice sheet) thule lower (thu_l; 570 m elevation) automated weather station located on tuto ramp. this station is representative of the warmest low-elevation conditions along the ice-sheet segment. we also assess the gc-net (greenland climate network) humboldt (hum) station in the ice-sheet interior, as representative of the coldest high-elevation conditions. the annual positive degree days assessed at thu_l during the 2010–2023 observational period ranges from 51 days in 2013 to 97 days in 2019 (fig. 9; fausto et al. 2021; how et al. 2022). during this period, the average melt onset at thu_l is 29 june, with the earliest melt onset recorded on 18 march 2013 and latest melt onset recorded on 9 august 2010. these positive degree days represent the cumulative number of degrees celsius that daily mean air temperatures are above 0°c during a single melt season. their cumulative values provide an indication of melt intensity. this thu_l positive degree day analysis suggests that melting conditions become typical after 1 june. traverse operations on the ice sheet should therefore conclude by the end of may, to avoid surface melt of the ice sheet. we also assess air temperatures at the hum weather station between 1995 and 2021 (fausto et al. 2021; vandecrux et al. 2023; how et al. 2022). at 1950 m elevation, hum is characteristic of the high-elevation firn plateau that comprises most of the traverse route. daily average air temperature at hum rarely a b fig. 8 view of the traverse. a: photograph of a relatively elevated portion of the tuto access road that remains snow-free in the spring (27 may 2023). photo: william colgan. b: photograph of a relatively depressed portion of the tuto access road where spring snow depth is >1 m (31 may 2023). photo: william colgan. the location of the photos is depicted in fig. 2. https://doi.org/10.34194/ft7sk529 https://geusbulletin.org/ colgan et al. 2025: geus bulletin 59. 8387. https://doi.org/10.34194/ft7sk529 10 of 17 geusbulletin.org exceeds 0°c, with the highest daily average air temperature being 1.6°c (fig. 9). surface melt is typically negligible in the high elevation ice-sheet interior of north-west greenland. instead, the primary temperature concern on the high-elevation firn plateau is exceeding the operational cold-limit of vehicles. at hum, the daily average air temperature during the april and may period ranges between –40°c and –12°c, with instances of the daily average air temperature below –40°c. the coldest ever recorded daily average temperature at hum is –51.0°c on 22 march 2001. the daily average temperature typically climbs above –30°c in mid-april. this likely serves as the earliest practical start for inland traverse operations. crevasses pose an additional hazard on the ice-sheet segment of the traverse. the emergence of large crevasses c. 40–45 km inland from tuto ramp in recent decades has effectively suspended traverses from tuto ramp since 2016 (fig. 10). by ascending the nuto ramp, any traverse effectively avoids this narrow region of high ice strain between ullip sermia (harald moltke bræ) and the tidewater glaciers draining into puisillip kangerlua (de dødes fjord). ice-surface velocities are readily available from offset tracking in satellite radar imagery (solgaard et  al. 2021). via the tuto approach, ice-surface velocities in the first c. 120 km of the traverse to sierra average 9.0 ± 6.8 m/yr (fig. 4). via the nuto approach they average only 5.8 ± 4.9 m/yr. while crevasses were fig. 9 daily air temperatures and cumulative positive degree days (pdd) for all available station years during the polar day period for a: thule lower (thu_l) and b: humboldt (hum) ice-sheet weather stations (fausto et al. 2021; vandecrux et al. 2023; how et al. 2022). thu_l is located at an elevation of 570 m at tuto ramp near pituffik. hum is located at an elevation of 1950 m in the high-elevation ice-sheet interior. the transparent range around the daily air temperature values represents an uncertainty of one standard deviation. https://doi.org/10.34194/ft7sk529 https://geusbulletin.org/ colgan et al. 2025: geus bulletin 59. 8387. https://doi.org/10.34194/ft7sk529 11 of 17 geusbulletin.org historically present in ablations areas along the inland tuto approach to sierra (meier et al. 1957), they have never been historically present in the nuto ablation area and approach to sierra (nobles 1960). crevasse traces, however, are currently clearly visible on nuto ramp (fig. 11). these crevasse traces are distinctive bands of dark ice that result from refrozen water within healed fractures. they can originate from several processes, including (1) recent cessation of local crevassing as ice flow becomes less dynamic, (2) ongoing thrust fracturing and annealing occurring locally at similar time scales and/or (3) advection of re-healed full-thickness crevasses from upstream (hambrey 1976; hambrey & müller 1978; hambrey & lawson 2000). we cannot speculate on the origin of the pronounced crevasse traces in the ablation area of nuto ramp today, but the presence of well-developed supraglacial runoff channels that consistently bisect these features indicate that they do not disrupt supraglacial hydrology by providing englacial access points. this suggests that they are not open features that can impend the safety or trafficability of ascending the ramp. although crevasses are not common within greenland’s high-elevation firn plateau, transverse and shear crevasses, both open and closed, of multi-metre widths have been documented above 1850 m elevation in a variety of different ice flow settings (mankoff et al. 2020; franke et  al. 2022; løkkegaard et  al. 2024). therefore, while crevasse hazard is generally minimised by selecting a route that minimises both absolute velocity and spatial gradients in velocity (i.e. strain rate), it is advisable to manually inspect satellite radar imagery of the entire high-elevation firn plateau route segment for crevasses. it is especially important to confirm the absence of crevasses in the vicinity of the inland ice stream onset that forms the petermann glacier. such an operational analysis could necessitate route deviations on the scale of 10s of kilometres. warming land science plan the ice-sheet ablation area in warming land, between sermeq konrad steffen to the west and sermeq niels reeh to the east, is the ultimate science goal of the 778 km traverse. there, a <2 km long transect spans the past 12 greenland interstadials and past 13 greenland stadials, allowing ice to be potentially dated and sampled back to c. 48 000 years before the present time (fig. 12; macgregor et al. 2020). this sampling transect starts at fig. 10 crevasse hazards along the nuto and tuto approaches. a: no crevasses are evident in the ablation area segment of the nuto approach, even in tandem-x radar imagery from april 2022 with a 6 m resolution. b: ablation area crevasse hazards along the tuto approach in visible spot 6 satellite imagery from july 2020 with a spatial resolution of 1.6 m (klimadatastyrelsen 2024b). the red lines in insets c–i denote crevasses that intersect the tuto approach. the black dashed line represents the traverse route via both nuto and tuto approaches. https://doi.org/10.34194/ft7sk529 https://geusbulletin.org/ colgan et al. 2025: geus bulletin 59. 8387. https://doi.org/10.34194/ft7sk529 12 of 17 geusbulletin.org the holocene boundary (81.16217°n and 53.55995°w) and proceeds down the glacier, towards the ice margin, to the discontinuity that appears to mark the start of stratigraphically disordered basal ice (81.17447°n and 53.49773°w). this sampling transect is orientated to minimise crossings of large supraglacial channels. even older ice is likely found between the gi-13 termination (c. 48 000 years ago) and the ice margin but it is more difficult to map the chronology from space. the surface expression of the stadial and interstadial ice bands are all at least >20 m wide, which potentially allows them to be located with the precision of a handheld gps. our primary science goal is bulk sampling of pre-holocene ancient ice. following previous studies that have bulk-sampled greenland ice for subsequent chemical analysis (reeh et al. 2002; schaefer et al. 2009), we expect to use electric chainsaws, appropriately prepared and cleaned, to reach a depth of >1 m below the weathering crust, from where ice samples will be recovered. ice samples will ideally be recovered from a variety of depths within each of the stadial and interstadial ice bands. rather than transporting the ice back for laboratory analysis, we anticipate extracting and concentrating the air bubbles from the ice, and only transporting this extracted air back for laboratory analysis. this requires a field-operable vacuum-melting tank that uses a noble gas for evacuation. while this is a sensitive apparatus, it has been successfully used in ice-sheet field camps (petrenko et al. 2006, 2008). beyond paleoclimatology and the bulk sampling of ancient atmospheres, there are numerous secondary science goals to investigate in the warming land icesheet ablation area. these include ice-penetrating radar surveys to reconstruct three-dimensional englacial layering of stadial and interstadial ice bands to inform iceflow models (mackay et al. 2014), direct measurements of the material properties of ice deposited during different stadials and interstadials (paterson 1991), the first systematic survey of greenland’s most promising outcrop for meteorites (whillans & cassidy 1983), assessing the current mass balance and future persistence of greenland’s oldest ice outcrop (gilbert et al. 2017) and lastly using englacial temperature profiles to infer geothermal heat flow in a poorly unconstrained area (colgan et al. 2022). there is also the possibility to undertake further opportunistic studies at automatic climate stations located along the traverse route (matoba et  al. 2018; vandecrux et al. 2023; fig. 3). we anticipate that the science traverse will need to transport c. 3000 kg of scientific and camping equipment to warming land and back, which would support a six-person team for an approximately 1-month period (colgan et  al. 2018). following a series of small daily transit camps during the inbound traverse, the team would establish a larger multi-week science camp in the middle of the sampling transect. a science camp established at the gi-8/gs-8 transition on the sampling transect (81.1698°n, 53.51036°w) would be within c. 600 m fig. 11 assessing potential hazards along the nuto ramp. (a): visible band spot 6 satellite image of 1.6 m spatial resolution of nuto ramp in july 2020 (klimadatastyrelsen 2024b). the black dashed line represents the traverse route. locations of inset images are shown. b: highlights crevasse traces intersected by supraglacial streams. c: highlights supraglacial streams only. https://doi.org/10.34194/ft7sk529 https://geusbulletin.org/ colgan et al. 2025: geus bulletin 59. 8387. https://doi.org/10.34194/ft7sk529 13 of 17 geusbulletin.org of all ice bands between gi-13 termination (c. 48  000 years ago) and the younger dryas (gi-1 onset; c. 12 000 years ago; fig. 12). similar to the nuto ramp ablation area, the presence of well-developed supraglacial runoff channels in high-resolution visible imagery indicates an absence of crevasses in the warming land ablation area. this should permit field scientists to move with relative ease along the sampling transect. the green transition traverses are an economical way to move heavy equipment and supplies across the ice sheet, while also allowing scientific activities to be performed along the traverse route (fig. 13; lever & weale 2011; hawley et al. 2014; lever et al. 2016; stockholm et al. 2021). for example, the usace estimated that the weasel m29c, with a tare weight of c. 2500 kg and a ground pressure of 2 psi (14 kpa), has a rolling resistance of just 1% in dry snow conditions characteristic of much of the ice-sheet interior (nuttall & finelli 1955; rula 1960). under saturated snow conditions, found in the ablation area during the melt season, the weasel’s rolling resistance can increase to 10%. generally, however, the relatively low carbon footprint of ice-sheet traverses, in comparison to charter aircraft-supported ice-sheet access, makes them increasingly attractive options for adopting more c b older boundaries 1 1 2 12 13 14 12 1110 11108 98 9 7 76 65.23 4 5.1 14 15 16 15 16 17 172 3 4 5.1 5.2 greenland stadial (gs-) greenland interstadial (gi-) 13 fig. 12 overview of the planned ice sampling transect in warming land adapted from macgregor et al. (2020). a: contrast-stretched worldview-2 multispectral image of the ice-sheet margin at warming land with 1.8 m spatial resolution from 19 august 2014. white dashed line indicates area covered by panel b. ‘b’ indicates the origin (0 m) of the science transect. b: 45º-rotated zoom-in on exposed surface layering with sampling transect (black line). colour shading of lines defined in panel a. c: dust concentration measured in the northgrip ice core (ruth et al. 2003). shaded polygons that link panel b with c illustrate potential stratigraphic correlations between surfaceand ice-core observed dustiness – blue shading indicates low dust concentrations and brown shading indicates high dust concentrations. d: relation between along-transect distance and apparent ice age (based on visual correlations; colour shading of dots defined in panel a). worldview image in a and b: copyright maxar inc. https://doi.org/10.34194/ft7sk529 https://geusbulletin.org/ colgan et al. 2025: geus bulletin 59. 8387. https://doi.org/10.34194/ft7sk529 14 of 17 geusbulletin.org sustainable ice-sheet operations as part of the green transition. the fuel consumption of ice-sheet traverse vehicles is highly variable. reported fuel consumption varies with vehicle type, payload, operating altitude, ice-sheet conditions and driver performance. we tabulate reported mileage and payloads for six vehicle types and estimate their carbon efficiency as vehicle mileage divided by vehicle payload (table 2). the carbon efficiency units of l/km/t reflect the amount of fuel required to move a payload of one tonne over a distance of 1 km. these efficiencies range across an order of magnitude, from 0.023 l/km/t for a piston bully 300 to 0.80 l/km/t for a weasel m29c. for comparison, a ski-equipped twin otter aircraft consumes c. 2 l/km of jet-a fuel and can deliver 300 kg of payload to warming land, which is at an effective range limit from qaanaaq. this yields a somewhat analogous carbon efficiency of 6.7 l/km/t, but this is not entirely comparable to the values listed in table 2 as it externalises fuel weight. generally, a ground traverse to warming land will use at least 90% less fuel per kilometre per tonne than typical aircraft-supported operations. traverse vehicles also offer the possibility to use bioor alternative fuels, fig. 13 photographs highlighting the light traverse vehicles itemised in table 2. a: snowmobile. photo: william colgan. b: weasel. photo: wikimedia commons. c: pisten bully. photo: patrick smith. d: tucker snocat. photo: patrick smith. e: arctic truck. photo: patrick smith. f: sherp. photo: henrik p.ø. lassen. https://doi.org/10.34194/ft7sk529 https://geusbulletin.org/ colgan et al. 2025: geus bulletin 59. 8387. https://doi.org/10.34194/ft7sk529 15 of 17 geusbulletin.org which are not presently an option for aircraft operations in greenland. for our specific warming land traverse application, we can calculate the direct carbon emissions associated with deploying c. 3000 kg with six persons to our intended science camp via twin otter versus sherp 1200n. for a twin otter scenario, each round-trip flight from qaanaaq to warming land to deliver 300 kg payload uses c. 2120 l of jet-a fuel. ten put-in flights and pull-out flights to transport the total camp payload would therefore require c. 42 400 l of jet-a fuel. for a sherp 1200n traverse, each vehicle would use c. 620 l of fuel for the return trip from pituffik to warming land and back to pituffik. we therefore assume a vehicle payload is approximately half fuel and half cargo. this would require five vehicles to transport the total science payload. these five vehicles would require a total of c. 3090 l of diesel fuel to complete the roundtrip traverse. in this specific application, the overland traverse scenario uses c. 7% of the carbon of the aircraft scenario. this difference in direct carbon emissions increases if the emissions associated with sea-lifting vehicles or flying aircraft to and from pituffik are also included. summary we evaluate the feasibility of an overland traverse from pituffik to warming land that provides a safe and efficient method to access greenland’s oldest ice outcrop. the 778 km route is designed to minimise variations in altitude and ice-surface velocities and strain rates, and mitigate potential sea-ice and crevasse hazards, as well as temperature risks. the route uses the historical nuto ramp to ascend the ice sheet, instead of the more crevassed tuto ramp. analysis of satellite imagery confirms that the primitive access road to nuto ramp is mostly drivable, with seven washouts in depressions that spring snowpack is expected to largely mitigate. the sea ice within uummannap kangerlua (wolstenholme fjord) has a relatively consistent seasonal pattern of thickness and extent. landfast sea ice with minimal ridging generally persists to the end of may. weather station data suggest that optimal conditions for snowpack firmness along the sea-ice and ice-sheet segments occur from mid-april to mid-may. on the ice sheet, during this period, there is both a low potential for extreme cold conditions as well as surface melting. annual vetting of the traverse route, including ice-penetrating radar surveys and shallow coring on the sea-ice segment and manual inspection of remotely sensed imagery for potential crevasse hazards on the ice sheet segment, is required for operational safety. these potential issues could result in minor route deviations to bypass transient hazards. accessing warming land via ground traverse can consume 90% less fuel than accessing the site through aircraft-supported operations. traverse vehicles also allow for the potential use of alternative fuels, which may further increase sustainability. by assessing logistical challenges, this study demonstrates a feasible, safe and low-impact method to access greenland’s oldest ice outcrop and other secondary science sites within the region. acknowledgements this report is a contribution of the traverse science planning group organised by geus. the traverse science planning group can be reached at traverse@geus.dk. this work supports the fieldwork activities of greenland climate network (gc-net) funded by the danish ministry of climate, energy and utilities. we thank erich osterberg (dartmouth college), andreas rønne stokholm (technical university of denmark), beata csatho (state university of new york) and h. jay zwally (retired) for sharing knowledge of some previous ice-sheet traverses. we thank henrik p.ø. lassen (arctic capacity/sherp denmark) for providing information about the sherp 1200n. we thank sune tamm (arctic trucks) for providing information about the arctic truck 6×6. we thank the three reviewers, scientific editor harold lovell and editor in chief catherine jex for their detailed comments, which substantially improved the paper. table 2 observed mileage and payloads, with calculated carbon efficiencies, of differing traverse vehicles under traverse conditions. the carbon efficiency units of l/km/t can be interpreted as the fuel required to move a payload of 1 tonne over a distance of 1 km. vehicles are shown in fig. 13. vehicle (fuel) mileage (l/km) payload* (kg) carbon efficiency (l/km/t) source pisten bully 300 (diesel) 0.80 34 090 0.023 patrick smith, per. comm. arctic trucks ford 6×6 (jet a1) 0.55 6100 0.090 sune tamm, per. comm. tucker snocat (diesel) 2.24 9070 0.25 lever & weale 2011 tucker snocat (diesel) 1.88 7270 0.26 patrick smith, per. comm. pisten bully 100 (diesel) 1.88 6820 0.28 patrick smith, per. comm. arctic truck toyota 6×6 (jet a1) 0.62 2200 0.28 sune tamm, per. comm. sherp n1200 (diesel) 0.40 1250 0.32 henrik lassen, per. comm. arctic truck toyota 6×6 (diesel) 0.62 910 0.68 patrick smith, per. comm. polaris 550 snowmobile (gasoline) 0.40 550 0.72 patrick smith, per. comm. weasel m29c (gasoline) 1.27 1590 0.80 benson 1955 *payload refers to the combined vehicle and sled loads, including passengers and fuel. https://doi.org/10.34194/ft7sk529 https://geusbulletin.org/ mailto:traverse@geus.dk colgan et al. 2025: geus bulletin 59. 8387. https://doi.org/10.34194/ft7sk529 16 of 17 geusbulletin.org additional information funding this work was supported by gc-net. author contributions conceptualisation: w.c., n.b.k., r.s.f. and c.s. analysis: d.f., g.l., a.l., e.b.n., p.h., a.r., a.s., r.s.f., k.k.k., j.a.m. and p.s. data curation: p.s. investigation: all authors. visualisation: w.c., d.f., g.l., a.l., p.h., a.r., a.s., k.k.k., j.m. and r.s.f. writing: w.c., n.b.k., d.f., g.l., a.l., p.h., a.r., a.s., j.a.m. and k.k.k. administration: a.p.a., w.c. and n.b.k. competing interests henrik lassen and sune tamm supply data on their own vehicles. additional files a shapefile of the pituffik to warming land traverse route described here is available in the online supplementary mate rial associated with this report (supplementary file s1; colgan et al. 2025). https:// 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https://doi.org/10.1016/j.quascirev.2005.09.005 https://doi.org/10.3189/002214308784886135 https://doi.org/10.3189/002214308784886135 https://doi.org/10.7910/dvn/3vdc4w https://doi.org/10.3189/172756402781817365 http://hdl.handle.net/11681/2137 http://hdl.handle.net/11681/2137 https://doi.org/10.1029/2002jd002376 https://doi.org/10.3189/002214309788816704 https://apps.dtic.mil/sti/tr/pdf/ad0745150.pdf https://doi.org/10.22008/fk2/lydwea https://doi.org/10.5194/essd-13-3491-2021 https://doi.org/10.5194/essd-13-3491-2021 https://doi.org/10.34194/geusb.v47.5369 https://doi.org/10.3189/s002214300000215x https://apps.dtic.mil/sti/tr/pdf/ad0263548.pdf https://apps.dtic.mil/sti/tr/pdf/ad0263548.pdf https://doi.org/10.5194/essd-15-5467-2023 https://doi.org/10.1126/science.222.4619.55 traverse route from pituffik to warming land, north greenland greenland’s oldest ice outcrop motorised ice-sheet traverses pittufik to warming land sea-ice segment primitive road segment ice-sheet segment warming land science plan the green transition summary additional information funding author contributions competing interests additional files references figures fig. 1 overview placing the traverse route assessed in this study in the context of past traverses fig. 2 overview of the study area. a: highlighting the tuto and nuto ramps, as well as place names i fig. 3 overview of the traverse route from pituffik to sermeq niels reeh in warming land, with ice-s fig. 4 feature of the traverse routes. a: ice surface velocity along the tuto and nuto traverse rout fig. 5 sea-ice conditions in the vicinity of the traverse routes. a: sea-ice development stage asse fig. 6 views to the traverse. a: oblique aerial photograph of land-fast ice within uummannap kanger fig. 7 map of a primitive road on the nunatarsuaq plateau between nunatarsuup nuua (kristiansen poi fig. 8 view of the traverse. a: photograph of a relatively elevated portion of the tuto access road fig. 9 daily air temperatures and cumulative positive degree days (pdd) for all available station y fig. 10 crevasse hazards along the nuto and tuto approaches. a: no crevasses are evident in the abl fig. 11 assessing potential hazards along the nuto ramp. (a): visible band spot 6 satellite image o fig. 12 overview of the planned ice sampling transect in warming land adapted from macgregor et al. fig. 13 photographs highlighting the light traverse vehicles itemised in table 2. a: snowmobile. ph tables table 1 acquisition dates of the available landsat 8 cloud-free imagery used for mapping of sea-ice table 2 observed mileage and payloads, with calculated carbon efficiencies, of differing traverse ve geological survey of denmark and greenland bulletin 15, 2008, 65-68 in recent years the arctic has come into focus for hydrocarbon exploration, and areas offshore both west and east greenland have been evaluated as promising frontier hydrocarbon provinces. seven hydrocarbon exploration and ex ploitation licenses were awarded in 2007–2008 offshore the disko–nuussuaq region (fig. 1), and two more have been awarded in the open-door region offshore south-western green land. in 2007, an extensive amount of new seismic and aeromagnetic data was acquired by the tgs-nopec geo physical company in the north-eastern baffin bay region. geophysical mapping has been initiated by the geological survey of denmark and greenland (geus) in the melville bugt region offshore north-west greenland (fig. 1) with the purpose of evaluating the hydrocarbon prospectivity. initial interpretation of seismic and gravity data suggests the presence of deep sedimentary basins separated by structural highs. the north-east baffin bay region, offshore greenland – a new frontier petroleum exploration region ulrik gregersen © geus, 2008. geological survey of denmark and greenland bulletin 15, 65–68. available at: www.geus.dk/publications/bull 65 fig. 1. gravity map (100 km filtered bouguer gravity data) of the study region, showing the seismic grid. the location of the seismic line shown in fig. 2 is indicated. large areas with marked high gravity values (red) often indicate structural highs, whereas areas with low values (blue) mostly indicate deep sedimentary basins. most basins and marked highs indicated by gravity data are confirmed by seismic data (fig. 2). the gravity data are public domain data from the danish national space center and have been filtered by geus. disko greenland svartenhuk halvø nussuaq umiivik-1 melville bugt baffin bay 500 km kanumas preference areas 76°n76°n 75°w75°w 70°w70°w 65°w 76°n 76°n 74°n74°n74°n 75°w 75°w75°w75°w 60°w70°w k i v i o q r i d g e k i v i o q r i d g ebaffin baybaffin bay melville bugt melville bugt 100 km greenland fig 2 kan-92 seismic survey bmp-00 seismic survey dw-02 seismic survey baffin bay melville bugt n k i v i o q r i d g e 66 geological information on source rock, reservoir rock and seal intervals from surrounding regions suggest that the melville bugt region is likely to have a significant petroleum potential. the study is based on public domain magnetic and gravity data, and all proprietary and public 2-d seismic data (fig. 1) acquired before 2003. seismic horizons from the ‘seismic basement’ to ‘base quaternary’ are being interpreted regionally. based on the seismic interpretation, a structural element map, depth-structure maps and isopach maps will be produced in order to assess the prospectivity of the melville bugt region. structures and basin development interpretations of structures and basin development are main ly based on seismic data supplemented by gravity and magnetic data. because no wells have been drilled in eastern baffin bay, interpretations of the geological development and petroleum prospectivity are based partly on correlations to wells farther south and partly on the onshore geology of the surrounding region. the mapping and regional analyses show deep basins and mainly north-north-west-trending major structural highs (figs 1, 2). the melville bugt region is proba2 t w ow ay t ra ve l t im e (s ec ) 4 6 w e kivioq basin melville bay ridge melville bay graben fig. 3 5 km mc tc me mm mc dr tc me mm dr mc fig. 2. seismic section bmp00-209, melville bugt region, eastern baffin bay, illustrating the structural elements (for location of line, see fig. 1). intervals below the ‘mid-cretaceous’ (mc) and just below and above the ‘top cretaceous’ (tc) horizons may, by analogy with surrounding regions, contain reservoir levels. structural closures are present at levels near the ‘mid-cretaceous’ and the ‘top cretaceous’ horizons in the basin and the graben areas. cretaceous structures are locally affected by palaeogene tectonic reactivation indicated by convex flexures above the tc horizon. the ‘mid-eocene’ (me) and the ‘mid-miocene’ (mm) horizons truncate structures, especially over the melville bay ridge, and the horizons are locally affected by neogene faulting. local u-shaped marked reflections with high amplitudes at the ‘mid-cretaceous’ horizon may indicate intrusions, e.g. in the kivioq basin (arrowed). the succession below the ‘deep reflection’ (dr) horizon loses seismic reflector continuity, and probably reflects mostly the crystalline basement. local scattered, low continuity reflections at the dr horizon may result from deeply buried proterozoic to palaeozoic sediments and volcanics, probably comparable to some of those known north of the study area. the frame marks the location of fig. 3. bly characterised by continental crust, whereas oceanic crust may be present farther west, in the central part of baffin bay (whittaker et al. 1997; oakey 2005). the highs and basins in the melville bugt region probably developed during several phases of mainly cretaceous to paleocene rifting and compressional strike-slip related palaeogene tectonism. the first rift phase probably occurred during early to midcretaceous time and resulted in mainly extensional faulting that created the two major structural highs in the region, the melville bay ridge and the kivioq ridge, and some normal faults in the adjacent grabens (fig. 2). the second rift phase occurred in late cretaceous and possibly in early paleocene time and resulted in normal faulting, partly reactivating earlier cretaceous faults. during the palaeogene, flexures and thrust faults developed related to compressional tectonics, which affected especially the central and northern parts of the region. the melville bay ridge was uplifted during the late cretaceous and palaeogene (fig. 2). the boundary fault east of the melville bay graben was active at least during the cretace ous and early palaeogene, giving way to major graben subsidence. in the graben, more than 12 km of sediments accumu lated that probably comprise mainly mesozoic and cenozoic deposits. in the lower part of the graben, some palaeozoic and proterozoic deposits similar to those known in north-west and north greenland may be present (peel & sønderholm 1991; dawes 1997). a more than 1.3 km thick succession of palaeogene basalts occurs in the southern part of the region and thickens southeastwards towards svartenhuk halvø. on svartenhuk halvø, nuussuaq and disko a thick palaeogene basalt succession is exposed (larsen & pulvertaft 2000; pedersen et al. 2006), and thick, tilted basalt successions are interpreted to extend out onto the adjacent offshore area (skaarup & pulvertaft 2007). offshore, the basalts occur in the upper part of some fault blocks that are on-lapped by subsequent deposits below the suggested mid-eocene horizon, indicating a mainly early eocene age of faulting and tilting of the blocks, which is also described by skaarup & pulvertaft (2007). deposition of thick sedimentary successions took place during the late cretaceous and in parts of the palaeogene. a major unconformity has been recognised at a mid-eocene level both in the baffin bay region (fig. 2) and offshore west greenland (dalhoff et al. 2003). neogene deposition is characterised by westwards and southwards prograding depositional systems, as also seen farther south offshore west greenland (dalhoff et al. 2003). prospectivity the melville bugt region in eastern baffin bay probably has a considerable hydrocarbon potential. major structures have been located at a suggested mid-cretaceous level (fig. 2), some with large structural closures at depths of c. 1–4 km. most of the structures have an attractive combination of size, short distance to potential mature source kitchens and a thick overburden. several potential reservoir rock levels may be present if cretaceous and palaeogene sand dominated units equivalent to those found farther south offshore and onshore west greenland are also present here (dalhoff et al. 2003; scherstén & sønderholm 2007). 67 w 2 km e 1.5 t w ow ay t ra ve l t im e (s ec ) 2 1 fig. 3. detail of the seismic section in fig. 2 showing abnormal seismic amplitude values (arrowed), interpreted as possible ‘bright spots’. 68 the deep basins of the melville bugt region contain thick successions of sedimentary rocks that may include prolific source rocks. several indications of an active hydrocarbon system with mature source rocks are known from canada and western greenland. in the arctic islands of canada these comprise exposed oil-prone source rocks of the kanguk formation (núñez-betelu et al. 1993) and the scott trough oil seep (maclean et al. 1981). in west greenland the presence of mature source rocks is proven by oil seeps onshore nuussuaq, disko and svartenhuk halvø (bojesen-koefoed et al. 1999, 2007). oil and gas have been encountered in boreholes on nuussuaq and svartenhuk halvø (e.g. wet gas in the umiivik-1 borehole; dam et al. 1998). ‘bright spots’ (fig. 3) and other hydrocarbon indicators related to structural closures or faults in the melville bugt region also suggest an active hydrocarbon system. a ‘bright spot’ is a type of hydrocarbon indicator that may point to, for example, gas in sands. potential mid-cretaceous source intervals may, as farther south, be in the ‘late oil window’ (gregersen et al. 2007), whereas the more deeply buried parts may be in the ‘gas window’. the distances from the source kitchens to the mapped structures are mostly less than 50 km. the cenozoic sediment cover generally exceeds 900 m over the cretaceous structures. the potential seal sections of late cretaceous and palaeogene ages are probably dominated by mudstones with local sands as known from outcrops and a few wells farther south (dalhoff et al. 2003; gregersen et al. 2007). the seismic data from the eastern baffin bay region (bmp-00 from the year 2000) have increased our knowledge about the main structures and faults and have facilitated the interpretation of the basins, but new data are needed to increase our understanding of the geology and prospectivity of the region. references bojesen-koefoed, j.a., christiansen, f.g., nytoft, h.p. & pedersen, a.k. 1999: oil seepage onshore west greenland: evidence of multiple source rocks and oil mixing. in: fleet, a.j. & boldy, s.a.r. (eds): petroleum geology of northwest europe: proceedings of the 5th conference, 305–314. london: geological society. bojesen-koefoed, j.a., bidstrup, t., christiansen, f.g., dalhoff, f., gre gersen, u., nytoft, h.p., nøhr-hansen, h., pedersen, a.k. & sønderholm, m. 2007: petroleum seepages at asuk, disko, west greenland – implications for regional petroleum exploration. journal of petroleum geology 30, 219–236. dalhoff, f., chalmers, j.a., gregersen, u., nøhr-hansen, h., rasmussen, j.a. & sheldon, e. 2003: mapping and facies analysis of paleocene – mid-eocene seismic sequences, offshore southern west greenland. marine and petroleum geology 20, 935–986. dam, g., nøhr-hansen, h., christiansen, f.g., bojesen-koefoed, j.a. & laier, t. 1998: the oldest marine cretaceous sediments in west greenland (umiivik-1 borehole) – record of the cenomanian–turonian anoxic event? geology of greenland survey bulletin 180, 128–137. dawes, p.r. 1997: the proterozoic thule supergroup, greenland and canada: history, lithostratigraphy and development. geology of greenland survey bulletin 174, 150 pp. gregersen, u., bidstrup, t., bojesen-koefoed, j.a., christiansen, f.g., dalhoff, f. & sønderholm, m. 2007: petroleum systems and structures offshore central west greenland: implications for hydrocarbon prospectivity. geological survey of denmark and greenland bulletin 13, 25–28. larsen, j.g. & pulvertaft, t.c.r. 2000: the structure of the creta ceous–palaeogene sedimentary-volcanic area of svartenhuk halvø, central west greenland. geology of greenland survey bulletin 188, 40 pp. maclean, b., falconer, r.k. & levy, e.m. 1981: geological, geophysical and chemical evidence for natural seepage of petroleum off the northeast coast of baffin island. bulletin of canadian petroleum geology 29, 75–95. núñez-betelu, l.k, riediger, c.l. & hills, l.v. 1993: rock-eval analysis of the cretaceous bastion ridge and kanguk formations, axel heiberg and ellesmere islands, canadian arctic archipelago. abstract (a-78), annual meeting of the geological association of canada, edmonton. oakey, g.n. 2005: cenozoic evolution and lithosphere dynamics of the baffin bay – nares strait region of arctic canada and greenland, 233 pp. unpublished ph.d. thesis, vrije universiteit, amsterdam, the netherlands. pedersen, a.k., larsen, l.m., pedersen, g.k. & dueholm, k.s. 2006: five slices through the nuussuaq basin, west greenland. geological survey of denmark and greenland bulletin 10, 53–56. peel, j.s. & sønderholm, m. (eds) 1991: sedimentary basins of north greenland. bulletin grønlands geologiske undersøgelse 160, 164 pp. scherstén, a. & sønderholm, m. 2007: provenance of cretaceous and paleocene sandstones in the west greenland basins based on detrital zircon dating. geological survey of denmark and greenland bulletin 13, 29–32. skaarup, n. & pulvertaft, t.c.r. 2007: aspects of the structure on the coast of the west greenland volcanic province revealed in seismic data. bulletin of the geological society of denmark 55, 65–80. whittaker, r.c., hamann, n.e. & pulvertaft, t.c.r. 1997: a new frontier province offshore northwest greenland; structure, basin development, and petroleum potential of the melville bay area. american association of petroleum geologists bulletin 81, 978–998. author’s address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ug@geus.dk larsen 2023: geus bulletin 54. 8347. https://doi.org/10.34194/geusb.v54.8347 1 of 2 preface preface rune berg-edland larsen* department of geoscience and petroleum, norwegian university of science and technology (ntnu), trondheim, norway *correspondence: rune.larsen@ntnu.no received: 02 feb 2023 accepted: 10 feb 2023 published: 27 jul 2023 keywords: igneous complex, magma chamber processes, mineralisation, platinum group elements, skaergaard intrusion abbreviations: au: gold pge: platinum group elements ntnu: norwegian university of science and technology geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: catherine jex (geus, denmark) reviewed by: none. preface to: rudashevsky et al. 2023 (this volume) funding: none. competing interests: none. additional files: none. in this special issue of geus bulletin, the many riddles regarding the platinum group elements and gold (pge-au) mineralisation of the east greenland skaergaard intrusion are untangled and discussed. the skaergard pge-au mineralisation, as defined in this study, embodies an enigmatic and rich ore formation that arguably could have been an economic resource, had it not been for its ice-locked position in central east greenland. the authors of this study (rudashevsky et al. 2023, this volume) characterise the systematic variability in the precious metal mineralogy from the contact towards the interior of the intrusion based on the analysis of more than 4000 individual pge-au grains. this variability is interpreted in the light of 90 years of research and over 1000 publications pertaining to magma chamber processes in the skaergaard intrusion. with such an impressive library of knowledge, on a comparatively simple magmatic system such as the skaergard intrusion, we should have discovered a few islands of truth in igneous petrology and ore-deposit formation. and indeed, we have. but we are also enriched with an evolving story, where answering one question only serves to raise three new questions. this study demonstrates the variability of pge-au phases throughout the ore-forming zone of the skaergaard intrusion. as previously observed, pge-au mineralisation in the central parts is divided into several layers over 30–40 m of the cumulus stratigraphy with increasing pd/pt ratios upwards, an au-rich upper part and a low sulphide content throughout all layers. close to the contact, the precious metal zonation is less pronounced, and it is significantly more sulphide rich. the pge mineralogy deviates significantly from the centre to the margin. these complex lateral and vertical variations cannot be explained by one genetic model but require an intricate combination of igneous processes including silicate-melt liquid immiscibility, sulphide-melt immiscibility, sulphide-melt resorptions, precious metal transport by volatile-rich fluids and, finally, the solidification rate of the cumulus mushes. for other well-preserved pge-au deposits throughout the world, we observe a great variation of ore-forming models. remarkably, most of these models may be applied to various parts of the skaergaard mineralisation. the authors suggest that the skaergaard intrusion preserves different steps in pge-au ore genesis, which in many other intrusions are obliterated by later https://doi.org/10.34194/geusb.v54.8347 https://orcid.org/0000-0002-0117-1106 mailto:rune.larsen@ntnu.no larsen 2023: geus bulletin 54. 8347. https://doi.org/10.34194/geusb.v54.8347 2 of 2 www.geusbul let in.org igneous events. therefore, the legacy of the skaergaard mineralisation is the preservation of igneous ore-forming events that may also precede the genesis of other pge-au deposits in the world. after the last conclusion, i guarantee that you will be confused and perhaps a bit triggered but hopefully also inspired and bursting with new questions on the genesis of pge-au deposits in mafic and ultramafic igneous complexes. in light of the recent study, you may even be encouraged to look at your favourite pge-au deposit with fresh eyes. this study beautifully demonstrates that turning the next page in the book of magma chamber processes is more important than seeing ‘the end’. rune b. larsen professor, economic geology, ntnu, norway reference rudashevsky, n.s., nielsen, t.f.d. & rudashevsky, v.n. 2023: the pge-au mineralisation of the skaergaard intrusion: precious metal minerals, petrography and ore genesis. geus bulletin 54 (this volume), 8306. https://doi.org/10.34194/v.54.8306 https://doi.org/10.34194/geusb.v54.8347 http://www.geusbulletin.org https://doi.org/10.34194/v.54.8306 geological survey of denmark and greenland bulletin 26, 2012, 29-32 29 early holocene sea-level changes in øresund, southern scandinavia ole bennike, martin skov andreasen, jørn bo jensen, matthias moros and nanna noe-nygaard the baltic sea and kattegat are connected via three straits: storebælt, lillebælt and øresund (fig. 1). øresund is the shallowest with a threshold around 7 m deep and increasing water depths to the north (fig. 2). in the early holocene, global sea-level rise led to reflooding of øresund. it started in northern øresund which was transformed into a fjord. however, so far the timing of the transgression has not been well determined, but sediment cores collected north of the threshold, at water depths of 12 to 20 m, and a new series of radiocarbon ages help to constrain this. as the relative sea level continued to rise, the threshold in øresund was also flooded, and øresund became a strait. in mid-holocene time, the relative sea level rose until it was 4–5 m higher than at present, and low-lying areas around øresund became small fjords. during the late holocene, the relative sea level fell again. part of the data set discussed here was presented by andreasen (2005). active glacier ice disappeared from the region around 17 cal. ka bp (calibrated kilo-years before present; houmarknielsen & kjær 2003). during deglaciation, the ice margin receded southwards and huge amounts of meltwater flowed northwards and formed deep holes and channels. shortly after the deglaciation of the region, marine waters inundated øresund and remains of arctic species such as polar cod (boreogadus saida ) and ringed seal (phoca hispida) have been found and dated to around 16 and 17.6 cal. ka bp (lagerlund & houmark-nielsen 1993). during this time, the relative sea level must have been high, reflecting the loading of the fennoscandian ice sheet during the last glacial maximum. somewhat later øresund played an important role in draining the baltic ice lake, and a major delta formed north of øresund. radiocarbon dating of marine mollusc shells from the delta has given ages between 16.3 and 11.8 cal. ka bp. following isostatic rebound the relative sea level fell, and major parts of øresund became dry land with forests, lakes and peat bogs. submarine lake and peat deposits from the younger dryas and the early holocene have been reported (jessen 1923). several now submarine former settlements have been reported from øresund (fischer 1993), which can be referred to the mesolithic based on artefacts and radiocarbon dating. methods sediment coring was carried out with a 10 cm diameter vibrocorer from r/v alexander von humboldt and r/v fig. 1. map of denmark and the surrounding area showing the location of øresund and other place names mentioned in the text. fig. 2. map of the øresund region showing the bathymetry and the location of vibrocores. t: threshold. 10°e 14°e 55°n 57°n sweden denmark germany kattegat storebælt baltic sea lillebælt femern bælt øresund fig. 2 skåne 50 km jylland fyn 56°n 10 km 55°n 13°30´e 0–5 5–10 10–15 15–20 20–25 25–30 30–35 35–40 40–45 45–50 >50 water depth (m) 246230 sjælland skåne 246240 348240 246250 246280 255150 348180 13°e t t t © 2012 geus. geological survey of denmark and greenland bulletin 26, 29–32. open access: www.geus.dk/publications/bull 3030 core/ n. lat. w. long. lab. no. material depth depth age ( c calibrated locality no. bct* (cm) b.s.l. (m) years bp) age (years bp)† * below core top. † calibration is according to the intcal09 data (terrestrial samples) and the marine09 data (marine samples). ‡ too young compared with the other dates from the core. $ fisher (1993). 246230 55°46.801´ 12°44.461´ aar-8284 mytilus edulis 105–115 21.2 8310 ± 70 246230 55°46.801´ 12°44.461´ aar-8285 mytilus edulis 542 25.5 9155 ± 70 246230 55°46.801´ 12°44.461´ aar-8286 phragmites australis 560–570 25.7 9445 ± 70 246240 55°46.440´ 12°44.392´ aar-8287 macoma balthica 120–130 18.0 9200 ± 70 246240 55°46.440´ 12°44.392´ aar-8288 macoma balthica 380–390 20.6 9380 ± 80 246250 55°57.811´ 12°36.327´ aar-8626 littorina littorea 58–60 16.4 8455 ± 35 246250 55°57.811´ 12°36.327´ aar-8627 littorina littorea 64–70 16.5 8565 ± 40 246250 55°57.811´ 12°36.327´ aar-8603 macoma balthica 270–280 18.7 9475 ± 35 246250 55°57.811´ 12°36.327´ aar-8601 woody plant roots 300–310 18.9 8820 ± 40 246280 55°56.104´ 12°36.696´ aar-8628 cerastoderma edulis 203–207 18.4 9270 ± 45 246280 55°56.104´ 12°36.696´ aar-8630 wood 215–223 18.6 8940 ± 40 246280 55°56.104´ 12°36.696´ aar-8604 phragmites australis 220–230 18.6 8975 ± 50 246280 55°56.104´ 12°36.696´ aar-8631 corylus avellana 352–356 19.9 8740 ± 35 255150 55°56.144´ 12°37.247´ aar-8632 scirpus spp. 190–200 20.3 8765 ± 40 348180 55°40.883´ 12°51.410´ poz-42496 cladium mariscus 365–370 15.7 8660 ± 50 348240 55°49.398´ 12°42.976´ poz-42497 phragmites australis 400 26.0 9240 ± 50 pilhaken aar-1225 corylus avellana 7.8 8120 ± 90 pilhaken t-10667 quercus sp. 7.8 7945 ± 75 svalerumpen k-6036 tree root 5.9 7680 ± 115 svalerumpen t-10665 pinus sylvestris 6.5 8225 ± 95 table 1. selected radiocarbon ages from øresund 14 8765–8981 9840–10 101 10 569–10 781 9907–10 129 10 147–10 333 9003–9109 9113–9301 10 253–10 372 9740–10 115 10 039–10 189 9941–10 193 9951–10 228 9606–9772‡ 9686–9887 9546–9662 10 298–10 498 8980–9259$ 8649–8977$ 8386–8586$ 9031–9300$ professor penck. coring positions were selected from highresolution shallow seismic profiles. the seismic equipment comprised a boomer, an x-star and a sediment echosounder. differential gps was used for navigation. the cores were split and described in the laboratory, and selected cores were subsampled for palaeoecological analyses. the samples were wet sieved on 0.4, 0.2 and 0.1 mm sieves, and the residue left on the sieves was analysed with a dissecting microscope. macrofossils were identified using a reference collection. remains of terrestrial plants and marine molluscs were used for accelerator mass spectrometry (ams) radiocarbon age determination (table 1). results most of the cores consisted of holocene marine clayey and silty mud, sometimes with marine sand in the upper part. one core (348180) contained rootlet-peat overlain by sediment with numerous fruits of cladium mariscus and shells of land gastropods. core 348240 contained sediment with in situ rootlets tentatively identified as phragmites australis. figure 3 shows an example of a macrofossil diagram. a total of 74 taxa were noted, but only 17 of them are shown in the simplified diagram. the sequence coarsens upwards; this is interpreted as an increased energy level as the sea level rose and the basin was transformed to a fjord. common remains of tree birch (betula sect. albae) and pine (pinus sylvetris, not shown) as well as rare remains of aspen (populus tremula) indicate that the shore was not far away. the lower part of the sequence is dominated by lacustrine taxa, such as the leach erpobdella sp. and the bryozoan cristatella mucedo. shells and head shields of freshwater fleas (cladocera) are also abundant. macrolimnophytes are represented by, for example, najas minor and zannichellia palustris. both can tolerate some brackish water influence, and the sediments contain rare remains of hydroidea and the ostracode cythoromorpha fuscata, implying a weak influence of brackish water. the presence of najas minor seeds indicates that summer temperatures were slightly higher than at present (bennike et al. 2001). this may be confirmed by the common presence of cladium mariscus fruits in core 348180, but this plant is calciphilous and its abundance in early holocene deposits may reflect that leaching of the soils was not as far advanced as today. in the middle part of the sequence a marked peak in scirpus spp. (mainly scirpus tabernaemontani) fruits is seen, and at about the same level phragmites australis seeds also show a maximum. this may imply erosion of reed beds as the sea level rose. at around 220 cm depth hydrobia ulvae and at around 180 cm depth ruppia sp., potamogeton pectinatus and littorina littorea appear, indicating increased salinity, followed by nucula nucleus, corbula gibba and other marine species that indicate a salinity similar to present values. 31 sea-level changes figure 4 shows a plot of the ages of radiocarbon dated samples against depth. we have included a few samples from archaeological sites (table 1; fisher 1993). the full black line shows relative sea-level changes during the mid-holocene according to christensen (2001), when a number of small transgressions and regressions occurred. a similar picture may have characterised the late holocene, but due to lack of data we have drawn a straight, stippled line for this time period. we have also drawn a straight line for the early holocene. we consider this unlikely, but the data do not allow us to draw a more elaborate line. the marine and terrestrial samples should fall respectively below and above the line. however, it is not possible to draw the line so that this is fulfilled. this means that some of the ages are not correct. several explanations for this can be offered. one factor to consider is differential isostatic rebound. the samples come from a fairly large region, but most of the isostatic rebound occurred prior to the holocene, and the marine limit only falls from c. 5.5 to c. 4 m from north to south (christensen 2001). hence we suggest that the depth of the dated samples should be moved by only 1–2 m, which does not change the picture. we consider two other factors more important. one is the marine reservoir effect, for which a value of 400 years was used, which is common for the region. however, large variations are seen from place to place, and in closed fjords it can be several hundred years greater than the regional value. variations of several hundred years may also take place over time (olsen et al. 2009). the other factor is that some mollusc species, such as macoma balthica used in this study, can take up old carbon from the sediment, and dating of such species may therefore also produce ages that are several hundred years older than contemporaneous terrestrial samples (mangerud et al. 2006). dating of terrestrial samples can also be problematic, and here we have excluded an age determination of a hazel (corylus avellana) nut fragment that appears to be too young. an explanation for this could be contamination by modern carbon in the laboratory. we consider the other ages fig. 4. tentative curve showing relative sea-level changes in øresund during the holocene, based on radiocarbon dated samples. the curve is compared with a curve based on numerical, geophysical modelling by lambeck (1999). ka: kilo-annum (one thousand years). fig. 3. strongly simplified macrofossil concentration diagram of core 255150 from øresund. the core consists of a lower unit of horizontally layered clay and silt and an upper unit of homogenous silty mud. a single sample has been dated to 9686–9887 cal. years bp (table 1). the red bars show remains that were not counted. marine brackish d ep th ( m ) 12 10 8 6 4 2 0 age (cal. ka bp) 0 terrestrial10 20 30 lambeck (1999; model) this study christensen (2001) b et ul a se ct . a lb ae ph ra gm ite s au st ra lis sc ir p us s p p. a lis m a p la nt ag oaq ua tic a n aj as m in or e rp ob de lla s p . o th ot ri ch ia s p . c ri st at el la m uc ed o z an ic he lli a p al us tr is po ta m og et on p ec tin at us r up p ia s p . c yt he ro m or p ha f us ca ta h yd ro id e a in d e t. h yd ro bi a ul va e li tt or in a lit to re a n uc ul a nu cl eu s c or bu la g ib ba 20 20 200 2 4 4 10 50 100 4 20 200 5 2 5 terrestrial telmatic lacustrine brackish marine 9.8 ka cl ay si ltl it h o lo gy d e p th ( cm ) 0 50 100 150 200 250 300 350 3232 of terrestrial material to be reliable and have drawn the curve so that the terrestrial samples fall above it. according to the curve, the relative sea level rose c. 25 m to the present level from 10 to 8 ka bp. we have compared the curve with a curve constructed by lambeck (1999), using numerical modelling, which shows a good fit (fig. 4). discussion and conclusions during the earliest holocene, large parts of øresund were dry land, but local lakes and bogs existed in depressions. the shore level of the southern baltic sea and kattegat reached a lowstand (björck 1995). as the water level in kattegat began to rise, a fjord with brackish water and limited water exchange formed in northern øresund. later, the ongoing eustatic sea-level rise led to increased salinity, and the fjord became larger. at the same time, the water level in the baltic basin also increased (jensen et al. 1999). the threshold in øresund was flooded between 9 and 8 ka, and øresund developed into a strait. the oldest dated marine shell from øresund gave an age of 10.3–10.4 cal. ka bp, but we suggest that this is somewhat too old, and marine water probably did not reach a level of around 25 m b.s.l. until c. 10 ka. however, at the entrance to øresund where the water depth is 35–40 m, marine waters may have begun to enter several millennia earlier according to lambeck’s model (fig. 4). in storebælt, the oldest dated marine shell gave an age of c. 8100 cal. years bp (bennike et al. 2004), in the lillebælt the oldest shell date is c. 7700 cal. years bp (bennike & jensen 2010) and in mecklenburger bucht, the oldest reported shell date is c. 7600 cal. years bp (bennike & jensen 1998). these differences partly reflect different threshold levels. however, freshwater drainage from the baltic basin through storebælt may also have inhibited marine bivalves from entering this strait for centuries or millennia. acknowledgements the captain and crew of the former r/v alexander von humboldt and r/v professor penck of the institut für ostseeforschung in warnemünde are thanked for their help during marine cruises. references andreasen, m.s. 2005: træk af øresunds udviklingshistorie gennem tidlig holocæn, 151 pp. unpublished cand. scient. thesis, københavns universitet, danmark. bennike, o. & jensen, j.b. 1998: lateand postglacial shore level changes in the southwestern baltic sea. bulletin of the geological society of denmark 45, 27–38. bennike, o. & jensen, j.b. 2010: postglacial, relative shore-level changes in lillebælt, denmark. geological survey of denmark and greenland bulletin 23, 37–40. bennike, o., jensen, j.b. & lemke, w. 2001: late quaternary records of najas spp. (najadaceae) from denmark and surroundings. review of palaeobotany and palynology 114, 259–267. bennike, o., jensen, j.b., lemke, w., kuijpers, a. & lomholt, s. 2004: lateand postglacial history of the great belt, denmark. boreas 33, 18–33. björck, s. 1995: a review of the history of the baltic sea, 13.0–8.0 ka bp. quaternary international 27, 19–40. christensen, c. 2001: kystbosættelse og havniveauændringer i stenalderen. in: jensen, o.l., sørensen, s.a. & hansen, k.m. (eds): danmarks jægerstenalder – status og perspektiver, 183–193. hørsholm: hørsholm egns museum. fischer, a. 1993: stenalderbopladser på bunden af øresund. del 1, det centrale øresund, 103 pp. copenhagen: miljøog energiministeriet, skovog naturstyrelsen. houmark-nielsen, m. & kjær, k.h. 2003: southwest scandinavia, 40–15 kyr bp: palaeogeography and environmental change. journal of quaternary science 18, 769–786. jensen, j.b., bennike, o., witkowski, a., lemke, w. & kuijpers, a. 1999: early holocene history of the southwestern baltic sea: the ancylus lake stage. boreas 29, 437–453. jessen, k. 1923: en undersøisk mose i rungsted havn og de senglaciale niveauforandringer i øresund. danmarks geologiske undersøgelse iv. række 1(18), 18 pp. lagerlund, e. & houmark-nielsen, m. 1993: timing and pattern of the last deglaciation in the kattegat region, southwest scandinavia. boreas 22, 337–347. lambeck, k. 1999: shoreline displacements in southern-central sweden and the evolution of the baltic sea since the last maximum glaciation. journal of the geological society (london) 156, 465–486. mangerud, j., bondevik, s., gulliksen, s., hufthammer, a.k. & høisæter, t. 2006: marine 14c reservoir ages for the 19th century whales and molluscs from the north atlantic. quaternary science reviews 25, 3228–3245. olsen, j., rasmussen, p. & heinemeier, j. 2009: holocene temporal and spatial variation in the radiocarbon reservoir age. boreas 38, 458–470. authors’ addresses o.b. & j.b.j., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: obe@geus.dk m.s.a. & n.n.n., department of geography and geology, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. m.m., the leibniz institute for baltic sea research, seestrasse 15, warnemünde, d-18119 rostock , germany. research article | short bennike et al. 2021: geus bulletin 47. 6530. https://doi.org/10.34194/geusb.v47.6530 1 of 7 late glacial and holocene shore-level changes in the aarhus bugt area, denmark ole bennike*1 , katrine juul andresen2 , peter moe astrup3 , jesper olsen4 , marit-solveig seidenkrantz2 1geological survey of denmark and greenland (geus), aarhus, denmark, 2department of geoscience and iclimate centre, aarhus university, denmark, 3moesgaard museum, aarhus, denmark, 4aarhus ams centre, department of physics and astronomy, aarhus university, denmark abstract we propose a new relative shore-level curve for the aarhus bugt area, an embayment in eastern jylland, denmark, based on a compilation of published and new radiocarbon ages of organic material. lakes existed in the area during the late glacial and early holocene. lake level rose gradually until the region was inundated by the sea at c. 9000 cal. years bp. the relative sea level reached a high stand at about 6000 cal. years bp, when the local relative sea level was c. 3 m above present-day mean sea level. the aarhus bugt area was inundated by the sea later than the limfjord area in northern jylland, but earlier than the lillebælt region in southern denmark. the shorelevel curves for these areas differ partly because the glacio-isostatic uplift was more pronounced in the limfjord area than farther south and partly because the northern regions were inundated by the sea earlier than the southern areas. introduction during the last glacial maximum, large parts of denmark were covered by the scandinavian ice sheet (houmark-nielsen et al. 2012). about 21 000 cal. years bp (before present, i.e. before 1950 ce), the ice sheet began to retreat partly because of melting and partly because of iceberg calving. as the colossal mass of glacier ice disappeared from the land areas, glacio-isostatic rebound began. uplift of the land is still ongoing, with highest uplift rates in the northern part of denmark and lowest in the southwestern part of the country (vestøl et al. 2019). concurrent with the glacio-isostatic rebound, global mean sea level also rose as large amounts of meltwater from the retreating ice sheets flowed into the world’s oceans. in total, sea level has risen about 125 m after the last glacial maximum (chapell & shackleton 1986; lambeck et al. 2014). the combination of land uplift and sea-level rise results in local and regional relative sea-level changes. these relative sea-level changes can be reconstructed by dating samples that can be related to a former high-tide level, so-called sealevel index points. however, in the inner danish waters, we do not have information on sea-level index points, instead we used dating of shells of marine *correspondence: obe@geus.dk received: 06 apr 2021 accepted: 28 jun 2021 published: 23 sept 2021 keywords: aarhus bugt, late glacial, shore-level changes, sea-level changes, quaternary abbreviations: aarams: aarhus ams centre ams: accelerator-mass spectrometry bp: before present geus bulletin is an open access, peerreviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: william colgan (geus) reviewed by: lars clemmensen (university of copenhagen, denmark), jason kirby (liverpool john moores university, uk) funding: see page 6 competing interests: none declared additional files: none provided https://doi.org/10.34194/geusb.v47.6530 https://orcid.org/0000-0002-5486-9946 https://orcid.org/0000-0001-8029-3234 https://orcid.org/0000-0002-7538-7014 https://orcid.org/0000-0002-4445-5520 https://orcid.org/0000-0002-1973-5969 mailto:obe@geus.dk https://creativecommons.org/licenses/by/4.0/deed.ast bennike et al. 2021: geus bulletin 47. 6530. https://doi.org/10.34194/geusb.v47.6530 2 of 7 www.geusbul let in.org gastropods or bivalves that lived below sea level, peat that accumulated in bogs above sea level or tree stumps or roots of land plants. archaeological finds from refuse layers can also provide knowledge of sea level at a given time. these limiting data can be used to reconstruct former sea level, but sea-level curves based on such data are less well constrained than sea-level curves based on index points. the aim of this paper is to present a shore-level curve for the aarhus bugt area (fig. 1), from where we have a fairly large number of radiocarbon ages from marine, lacustrine and terrestrial deposits (table 1). we have compiled 32 ages and propose a new curve for the late glacial and holocene relative shore-level changes in the area. we use the term shore-level change rather than sea-level change because we have constructed both lake-level and sea-level changes. the shore-level curve for aarhus bugt fills a knowledge gap on shore-level changes in denmark and adds to a growing number of shore-level curves from the region (e.g. bennike & jensen 2011; clemmensen et al. 2012, 2018; bennike et al. 2012, 2019; hede et al. 2015; sander et al. 2016). material and methods new ages used to reconstruct shore-level changes come partly from two vibrocores (502052 and 502017-1) collected in relation to mapping of sand and gravel resources and from a gravity core (au18-mg-09g) collected during a student cruise with the aarhus university research vessel aurora in 2018. core positions were selected based on shallow seismic data and sub-bottom profiles collected during the cruise. we also include new ages from sediment cores retrieved from brabrand sø, a lake that was formerly a fjord. in addition, ages from published literature concerning archaeological excavations on land (andersen & liversage 1994; heinemeier & rud 1999, 2001; kveiborg 2014), marine archaeological investigations in kalø vig (fischer & hansen 2005; astrup 2018), geological studies of aarhus bugt (jensen & bennike 2009; rasmussen et al. 2020) and an age of a pine (pinus sylvestris) stump that was found during deepening of the harbour at aarhus are included (heinemeier & rud 2000). the location of the dated samples appears in fig. 1. we estimate that the elevation uncertainty is up to ± 0.25 m. the material for radiocarbon dating has been dried and submitted to a variety of laboratories; most hjelm aarhus harbour pustervig åby renseanlæg 502017 502052m1/gc 174 gc 160 09g brabrand sø kalø vig 1-2 5 km djursland aarhus bugt ebeltoft lystrup enge 10°e 14°e 57°n sweden kattegat baltic sea jylland fyn 50 km sjælland 10.5°e 56°n lf lb fig. 1 map of the aarhus bugt area, showing localities with radiocarbon-dated samples. lf: limfjorden; lb: lille bælt (inset map). https://doi.org/10.34194/geusb.v47.6530 http://www.geusbulletin.org bennike et al. 2021: geus bulletin 47. 6530. https://doi.org/10.34194/geusb.v47.6530 3 of 7 www.geusbul let in.org table 1 selected radiocarbon ages from the aarhus bugt area, denmark. core/site latitude (°n) longitude (°e) laboratory number material elevation (cm) age (14c a bp)1 cal. age (a bp)2 ref. 502017-1 56.033 10.326 ua-57753 cornus sanguinea –1120 7972 ± 37 8648–8993 a 502052 56.112 10.478 aar-29102 menyanthes trifoliata –3318 10 007 ± 36 11 280–11 697 a 502052 56.112 10.478 aar-29103 b.nana+dryas+s.pol. –3420 10 158 ± 54 11 405–11 971 a gc 160 56.117 10.433 poz-7848 cerastoderma edule –2580 8290 ± 40 8628–9002 b gc 174 56.117 10.35 poz-10516 mytilus edulis –1640 7870 ± 50 8150–8477 b gc 174 56.117 10.35 poz-10517 cerastoderma lamarcki –1740 8240 ± 50 8567–8976 b m1 56.117 10.35 aar-16263 cerastoderma edule –2616 8349 ± 45 8690–9107 c m1 56.117 10.35 uba-19004 corylus avellana –2640 8565 ± 43 9473–9658 c m1 56.117 10.35 aar-18772 deciduous leaf fragment –2660 8432 ± 32 9328–9530 c m1 56.117 10.35 aar-18773 betula, meny, schoeno –2695 8910 ± 34 9906–10 183 c aarhus havn 56.147 10.24 aar-4859 pinus sylvestris stump –1350 8200 ± 70 9003–9406 d pustervig 56.158 10.208 aar-4161 littorina –36 4885 ± 50 4943–5366 e åby rense. 56.152 10.178 aar-12828 cerastoderma sp. –100 5179 ± 39 5325–5648 f brabrand sø 56.146 10.109 aar-30708 ostrea edulis –710 6260 ± 31 6495–6836 a brabrand sø 56.146 10.109 aar-30707 ostrea edulis –865 5862 ± 34 6079–6395 a brabrand sø 56.146 10.109 aar-30706 mytilus edulis –877 8512 ± 40 8967–9310 a brabrand sø 56.146 10.109 aar-30705 twig –890 9249 ± 40 10 262–10 560 a brabrand sø 56.146 10.109 aar-30704 populus tremula –898 9377 ± 37 10 500–10 702 a brabrand sø 56.143 10.09 aar-30093 schonoplectus lacustris –58 2831 ± 32 2853–3058 a brabrand sø 56.143 10.09 aar-30092 cerastoderma –62 5428 ± 33 5602–5909 a brabrand sø 56.143 10.09 aar-30091 littorina littorea –216 7274 ± 36 7564–7848 a brabrand sø 56.143 10.09 aar-30090 in situ root, woody plant –224 6895 ± 39 7624–7834 a kalø vig 1 56.219 10.383 aar-8413 in situ quercus root –670 7690 ± 45 8394–8586 g kalø vig 2 56.219 10.383 aar-27412 tree stump –750 7813 ± 75 8417–8977 h 09g 56.236 10.418 aar-30088 twig –1363 8219 ± 46 9022–9399 a 09g 56.236 10.418 aar-30089 twig –1359 8256 ± 41 9032–9412 a hjelm 56.135 10.8 aar-5486 littorina littorea +380 5525 ± 55 5674–6083 i lystrup enge 56.223 10.225 k-4053 corylus branches –85 6210 ± 105 6799–7411 j lystrup enge 56.223 10.225 k-5730 populus dugout boat –35 6110 ± 100 6741–7252 j lystrup enge 56.223 10.225 k-6012 tilia dugout boat –81 6550 ± 105 7259–7614 j lystrup enge 56.223 10.225 k-6335 quercus tree trunk –25 5450 ± 100 5950–6440 j lystrup enge 56.223 10.225 k-6397 quercus tree trunk –48 6570 ± 100 7273–7613 j 1 ages in conventional radiocarbon years bp (before present = 1950; stuiver & polach (1977)); 2 calibration to calendar years bp (2 sigma) is according to the intcal20 or marine20 data (reimer et al. 2020; heaton et al. 2020). ref.: references. a: this study, b: jensen & bennike (2009), c: rasmussen et al. (2020), d: heinemeier & rud (2000), e: heinemeier & rud (1999), f: kveiborg (2014), g: fischer & hansen (2005), h: astrup (2018), i: heinemeier & rud (2001), j: andersen & liversage (1994). samples, however, have been dated at the aarhus ams centre (aarams; marked aar in table 1). these are partly remains of land plants and shells from marine molluscs. most of the age determinations were performed by accelerator-mass spectrometry (ams) by measuring the ratio of 14c to 12c atoms (olsen et al. 2009), but ages marked k in table 1 are conventional 14c ages. the ages are stated in conventional radiocarbon years bp and corrected for isotope fractionation by normalising to a δ13c value of –25‰ vpdb (stuiver & polach 1977). the radiocarbon ages are calibrated to calendar years before now using the calib version 8.2 program (stuiver et al. 2021). for marine samples, we used the marine calibration curve marine20, and for terrestrial samples, we used the intcal20 curve. for marine samples, we used a reservoir age of 400 years (i.e. δr = –150 years). both the new ages and previously published ages have been (re)calibrated for this study. sediments and macrofossils the sediments in the cores from aarhus bugt area encompass till deposits from the last glaciation (weichselian), late glacial lacustrine clay, holocene non-marine deposits and, finally, holocene marine sediments. late glacial fossiliferous (terrestrial and lacustrine plants and invertebrates) deposits were found in vibrocore 502052, which was 11.6 m long and collected at a water depth of 29.1 m (fig. 1 and table 1). the late glacial flora comprised the woody plants betula nana, dryas octopetala, salix polaris and empetrum nigrum. macrolimnophytes were represented by menyanthes trifoliata, potamogeton filiformis, p. perfoliatus, p. natans, callitriche hermaphroditica and chara sp. invertebrates comprised the leach erpobdella sp., the ostracods cytherissa lacustris, limnocythere sp., candona sp., the gastropods valvata cristata and v. piscinalis, the bivalve pisidium sp. and the bryozoan cristatella mucedo. https://doi.org/10.34194/geusb.v47.6530 http://www.geusbulletin.org bennike et al. 2021: geus bulletin 47. 6530. https://doi.org/10.34194/geusb.v47.6530 4 of 7 www.geusbul let in.org a subsample of betula nana, dryas octopetala and salix polaris remains was dated to the younger dryas, a cold period at the end of the weichselian (table 1). this is in accordance with the flora and fauna, which are typical of younger dryas deposits in the region (bennike et al. 2004). the terrestrial plants indicate a tree-less, tundra-like landscape characterised by dwarf shrub heaths. the early holocene terrestrial flora from aarhus bugt included the trees betula sect. albae (tree birch), populus tremula (aspen), pinus sylvestris (pine) and alnus glutinosa (alder), indicating a landscape with open forests. vibrocore 502017-1 shows an example of a succession with clayey till, peat, lacustrine gyttja, marine mud and, finally, marine sand and gravel (fig. 2). the lower part of the peat is dominated by stems and leaves of the brown moss scorpidium, whereas the upper part of the peat is dominated by twigs, indicating that the former bog was overgrown by trees or bushes. woody plants are represented by alnus glutinosa, betula sect. albae and cornus sanguinea. a fruit stone of the latter, which came from the upper part of the peat, was dated to c. 8850 cal. years bp (fig. 2). the lacustrine gyttja is dominated by vegetative remains of phragmites; it also contains numerous shells of lacustrine cladocerans. the marine mud contains a mollusc fauna that indicate lowered salinity and the ostracod cyprideis torosa, which is typical of environments with low and strongly fluctuating salinities (frenzel et al. 2012; pint et al. 2012). the submerged macrolimnophyte ruppia indicates shallow water. the marine mud is overlain by marine sand and gravel, presumably reflecting an increasing energy level as the sea level rose. shore-level changes in fig. 3, we have plotted ages against elevation for marine, lacustrine and terrestrial deposits. based on these ages, we suggest a curve that shows the development of the shore level in the aarhus bugt region over the last 12 000 years. the dated material comes from different sources, and it is difficult to quantify the vertical error on the samples, but we suggest an error of ± 0.5 m. peat can be compacted significantly when covered by sand and many metres of seawater, which will lower the deposit (baeteman et al. 2012). the curve is, thus, not well constrained, and we only indicate a likely development with a dashed line. the oldest age based on remains of dwarf shrubs from vibrocore 502052 yielded an age of 11 405–11 971 cal. years bp (table 1). the lithology and fossil content of the core shows that at this time, there were lakes in the deep parts of aarhus bugt. the water level in the lakes rose in the following period, and the lakes became larger. in the early holocene, peat bogs were probably widespread in wet areas of the aarhus bugt area, whereas more dry areas were forested. based on the radiocarbon ages, it appears that the sea began to inundate the aarhus bugt area about 9000 years ago (fig. 3), as also concluded by rasmussen et al. (2020). the marine inundation occurred because the rising global sea level surpassed the local glacio-isostatic land uplift of the area. at c. 9000 cal. years bp, global mean sea level was approximately 15 m lower than today (lambeck et al. 2014). initially, mixing of freshwater and seawater in the littoral zone created brackish conditions, as evidenced by the occurrence of low salinity species associated with shallow-water conditions, such as the bivalves cerastoderma sp. and mytilus edulis, the gastropods littorina littorea and hydrobia sp. and the ostracod cyprideis torosa. the relative sea-level rose until about 6000 cal. years ago, when it reached its maximum (fig. 3). the timing of the sea-level maximum is constrained by dating of a shell of littorina from a raised beach ridge on the island of hjelm. the marine limit on the island is c. 3 m above mean sea level, but raised beaches occur up to 5.3 m (mertz 1924). the shell was found at an elevation of 3.8 m, and the beach was probably deposited during a storm. the shell yielded an age of c. 5900 cal. years bp (table 1; heinemeier & rud 2001). at aarhus, the marine 1 2 3 4 d ep th (m ) 0 mud cl ay si lt vf sand f m c vc gravel gr an pe bblithology homogenous sand cerastoderma fine sand, cerastoderma, littorina, mytilus cerastoderma, mytilus, hydrobia, ruppia, cyprideis homogenenous marine mud phragmites, cladocera homogeneous lacustrine gyttja peat clayey till cornus: 8850 cal. years bp twigs, alnus homogenous sand and gravel scorpidium, carex rostrata fig. 2 sedimentological log of vibrocore 502017-1, collected at 56.033°n, 10.326°e, at a water depth of 9.1 m. the presence of peat with remains of alnus overlain by lacustrine gyttja and marine mud shows that the area has been transgressed by the sea. a fruit stone of cornus sanguinea was dated to c. 8850 cal. years bp (table 1). https://doi.org/10.34194/geusb.v47.6530 http://www.geusbulletin.org bennike et al. 2021: geus bulletin 47. 6530. https://doi.org/10.34194/geusb.v47.6530 5 of 7 www.geusbul let in.org limit is c. 2.5 m, at ebeltoft, it is c. 3.5 m and in the northeastern part of djursland, it is c. 5 m above present levels (mertz 1924). over the last 6000 years, global sea levels have been largely stable (e.g. lambeck et al. 2014), whilst in the aarhus bugt area, this period is marked by land uplift out-pacing the rate of sea-level rise resulting in a fall in the relative sea level. we have indicated a steady decline until today, but this part of the curve is poorly constrained with data and very uncertain (fig. 3). dating of a mytilus edulis (blue mussel) shell from marine deposits in lake brabrand sø gave a surprisingly old age (fig. 3). this indicates that the reservoir age was more than 400 years in the early stages of the fjord. it is also seen that two ages from lystrup enge appear to be too old. these ages come from samples deposited in shallow water close to the former seashore. comparisons with other shore-level curves from denmark (bennike & jensen 2011; bennike et al. 2019) show similar trends to the curve from the aarhus bugt area (fig. 4). however, marine waters inundated the western limfjord earlier than aarhus bugt, which, in turn, was inundated earlier than southern lillebælt. raised marine deposits are found up to 5 m above present sea level in the western limfjord area, whereas raised marine deposits are not found in southern lillebælt, where the rate of sea-level rise surpassed the glacio-isostatic uplift. conclusions during the younger dryas, most of aarhus bugt was dry land with dwarf shrub heaths, but small lakes existed locally. in the earliest holocene, most of aarhus bugt was dry land, but lakes soon filled the deeper parts of the area. the lakes expanded in size and shore-level rose. during this period, the trees betula sect. albae, populus tremula, pinus sylvestris and alnus glutinosa immigrated to el ev at io n (m ) 0 –10 +5 –20 –30 12 000 10 000 8000 6000 4000 2000 0 age (cal. years bp) lake-level changes relative sea-level changes marine deposit lake deposit terrestrial hjelm brabrand sø lystrup enge fig. 3 radiocarbon ages from the aarhus bugt area, plotted against elevation. the length of the bars represents the uncertainty range in the calibrated age (table 1). the dashed curve shows our best estimate of the relative shore-level changes from c. 12 000 cal. years bp until today, with the shift from green to blue colour, indicating the time of the marine inundation of the aarhus bugt area. limfjorden aarhus bugt southern lillebælt marine brackish lacustrine 12 000 10 000 8000 6000 4000 2000 0 age (cal. years bp) 0 –20 –30 –10 e le va tio n (m ) fig. 4 comparison of shore-level curves for the western limfjord area (bennike et al. 2019), the aarhus bugt (this study) area and southern lillebælt region (bennike & jensen 2011). https://doi.org/10.34194/geusb.v47.6530 http://www.geusbulletin.org bennike et al. 2021: geus bulletin 47. 6530. https://doi.org/10.34194/geusb.v47.6530 6 of 7 www.geusbul let in.org the region forming open forests. rising global sea levels resulted in a marine inundation of the deepest parts of aarhus bugt at about 9000 cal. years bp, and the relative sea level rose gradually during the following millennia and reached a high stand at c. 6000 cal. years bp, as documented by raised beach ridges on hjelm island. we propose that the relative sea level fell gradually during the late holocene due to gradual glacio-isostatic rebound, but the timing is not yet fully constrained. acknowledgements we thank the captain (torben vang) and crew of r/v aurora for excellent help during the marine cruise and also thank the students on the course for their diligent work. uffe rasmussen and hans skov  from moesgaard museum are thanked for information on samples from hjelm, aaby renseanlæg and pustervig in aarhus. the shore-level curve was constructed as part of the project ‘kystzonens geodynamik i nationalpark mols bjerge’ (geocenter danmark 2021). we also thank the project group, bent odgaard and jens reddersen,  for good discussions. journal referees lars b. clemmensen and jason kirby provided constructive comments to the manuscript. funding statement this study was supported by geocenter denmark and national park mols bjerge. we also acknowledge funding through the danish council for independent research (grant nos. 7014-00113b (g-ice) and 0135-00165b (greenshelf) to mss), and the project has also received funding from the european union’s horizon 2020 research and innovation program under grant agreement no. 869383 (ecotip; 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https://doi.org/10.1016/j.csr.2008.08.007 https://doi.org/10.1016/j.csr.2008.08.007 https://doi.org/10.1073/pnas.1411762111 https://doi.org/10.34194/raekke2.v41.6827 https://doi.org/10.1111/j.1502-3885.2009.00088.x https://doi.org/10.1002/iroh.201111495 https://doi.org/10.1111/bor.12408 https://doi.org/10.1017/rdc.2020.41 https://doi.org/10.1017/rdc.2020.41 bennike et al. 2021: geus bulletin 47. 6530. https://doi.org/10.34194/geusb.v47.6530 7 of 7 www.geusbul let in.org lagoons and beach ridges as complementary sedimentary archives for the reconstruction of holocene relative sea-level changes. terra nova 28(1), 43–49. https://doi.org/10.1111/ter.12187 stuiver, m. & polach, h.a. 1977: discussion of reporting 14c data. radiocarbon 19(3), 355–363. https://doi.org/10.1017/s0033822200003672 stuiver, m., reimer, p.j. & reimer, r.w. 2021: calib 8.2 [www program]. http://calib.org (accessed june 2021) vestøl, o., ågren, j., steffen, h., kierulf, h. & tarasov, l. 2019: nkg2016lu: a new land uplift model for fennoscandia and the baltic region. journal of geodesy 93, 1759–1779. https://doi.org/10.1007/s00190-019-01280-8 https://doi.org/10.34194/geusb.v47.6530 http://www.geusbulletin.org https://doi.org/10.1111/ter.12187 https://doi.org/10.1017/s0033822200003672 http://calib.org https://doi.org/10.1007/s00190-019-01280-8 late glacial and holocene shore-level changes in the aarhus bugt area, denmark abstract introduction material and methods sediments and macrofossils shore-level changes conclusions acknowledgements funding statement author contributions references figures fig. 1. map of the aarhus bugt area, showing localities with radiocarbon-dated samples. lf: limfjord fig. 2. sedimentological log of vibrocore 502017-1, collected at 56.033°n, 10.326°e, at a water dept fig. 3. radiocarbon ages from the aarhus bugt area, plotted against elevation. the length of the bar fig. 4. comparison of shore-level curves for the western limfjord area (bennike et al. 2019), the aa table table 1. selected radiocarbon ages from the aarhus bugt area, denmark. nielsen et al. 2024: geus bulletin 57. 8372. https://doi.org/10.34194/geusb.v57.8372 1 of 10 research article applying geotechnical borehole databases in the search for interglacial deposits in denmark emil schou nielsen1* , trine kellberg nielsen1,2 , søren munch kristiansen3 1school of culture and society – department of archaeology and heritage studies, aarhus university, højbjerg, denmark, 2moesgaard museum, højbjerg, denmark, 3department of geoscience, aarhus university, aarhus, denmark abstract geotechnical investigations conducted in preparation for infrastructure development provide high-quality borehole data in standardised digital formats. in denmark, such geotechnical borehole data are not required to be reported to the national well database (jupiter) and are mainly archived in privately owned databases. accessible interglacial and interstadial terrestrial deposits are rare in denmark, and these borehole data have the potential to identify interglacial and interstadial deposits, with significant implications for ongoing palaeoclimate, palaeoenvironmental and archaeological research. in this study, we compiled data from six major geotechnical companies, resulting in a database with over 550 000 boreholes. from this database, we identified 1850 boreholes containing samples associated with interglacial and interstadial ages. through extensive filtering for well-documented lacustrine or palustrine deposits, we selected 161 boreholes and referenced them to 39 different geographical occurrences. of these 39 occurrences, 36 were either new terrestrial deposits or provided substantial new records to known interglacial and interstadial sites. our findings demonstrate that access to these privately owned geotechnical borehole data can be a valuable resource for identifying rare near-surface geological deposits, allowing the discovery of several new pleistocene sedimentary archives that warrant further investigation. *correspondence: emsn@cas.au.dk received: 20 feb 2024 revised: 17 june 2024 accepted: 13 aug 2024 published: 30 sep 2024 keywords: palaeoclimate, palaeoenvironment, lacustrine deposits, pleistocene, stratigraphy abbreviations: adna: ancient dna b.s.: below surface geus: geological survey of denmark and greenland igis: interglacial and interstadial mis: marine isotope stage pg: post-glacial sedadna: sedimentary ancient dna geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: kristian svennevig (geus, denmark) reviewed by: ole bennike (geus, denmark), joakim stiel korshøj (geo, denmark) and one anonymous reviewer funding: see page 9 competing interests: see page 9 additional files: see page 9 1 introduction over the past 50 years, there has been increasing interest in understanding long-term climatic variations using large-scale geological archives, such as ice sheets, deep-marine sediment beds and loess deposits. as our understanding of palaeoenvironment advances, identifying more high-resolution palaeoenvironmental and palaeoclimate archives is crucial for contemporary efforts to understand past and future climate change and human-environmental interactions. these advances offer new relevance to small-scale terrestrial geological archives, which capture regional and local variations at a high resolution (moreno et al. 2014). interglacial terrestrial deposits from small lake basins have been used since the early 20th century to establish one of the first terrestrial quaternary stratigraphies in northern europe (jessen & milthers 1928; andersen 1965). these deposits became renowned for well-preserved plant macrofossils and pollen, but besides attempts to constrain the weichselian ice extent (e.g. hansen 1976), sporadic discoveries (e.g. gormsen & hansen 1980; kronborg & odgaard 2004; bennike et al. 2019) and a few dedicated studies (kuneš et  al. 2013), interglacial terrestrial deposits have received limited attention in recent decades. advances in dating methods, biomarker analyses and sedimentary ancient dna (sedadna; parducci et al. 2017; dussex et al. 2021) provide promising avenues for gaining new insights into interglacial terrestrial deposits. these advances offer an unprecedented opportunity for a detailed understanding of past ecosystems, climate and glacial erosion history. https://doi.org/10.34194/geusb.v57.8372 https://orcid.org/0009-0005-4653-6956 https://orcid.org/0000-0001-6481-5661 https://orcid.org/0000-0003-3128-4061 mailto:emsn@cas.au.dk https://creativecommons.org/licenses/by/4.0/deed.ast nielsen et al. 2024: geus bulletin 57. 8372. https://doi.org/10.34194/geusb.v57.8372 2 of 10 geusbulletin.org renewed interest in danish interglacial deposits has recently emerged in the field of archaeology, as palaeolithic research has attempted to establish the limits of past human northward expansion (nielsen et al. 2019). in this context, the focus is primarily on the last interglacial stage and the following interstadials, namely the eemian interglacial (marine isotope stage [mis], mis 5e) and the two following interstadials, brørup (mis 5c) and odderade (mis 5a). these periods offer favourable conditions owing to their relatively young age and minimal glacial disturbance, providing the best opportunities for preservation and evidence of past human presence. on the north european plain in northern germany, eemian and early weichselian open-air neanderthal sites have been discovered in lacustrine settings at schalkholz (arnold 1978), lichtenberg (weiss et al. 2022) and neumark-nord (pop et al. 2015; roebroeks et al. 2021). at the latter site, a large assemblage of processed megafauna remains was excavated. earlier occupation of the north european plain is known from the site of schöningen (urban et  al. 2023b), which is correlated with mis 9 (urban et  al. 2023a). the schöningen site holds the largest assemblage of wooden hunting spears currently known from the pleistocene. if the aim is to locate new and well-preserved archaeological sites, lake shores offer a suitable habitat for early humans. adjacent lake sediments may also provide a dateable context for archaeological remains, such as lithics, and hold climate and environmental proxies. to date, the search for such terrestrial interglacial deposits in northern europe has been opportunistic, using coal and gravel pits, larger infrastructure projects or natural exposures to study interglacial deposits. in the search for traces of neanderthals who may have occupied the region of southern scandinavia during mis 5e (yaworsky et  al. 2024), this approach has so far been unsuccessful in denmark (nielsen & riede 2018), and a systematic method for locating interglacial and interstadial (igis) deposits is needed. the danish national well database, jupiter, is a readily available and accessible resource of borehole data in denmark maintained by the geological survey of denmark and greenland (geus). the jupiter database has extensive spatial and temporal coverage; however, it has limitations owing to the heterogeneity of the well data and a lack of context, as it consists primarily of single-point records. notably, no systematic study has been conducted to identify interglacial deposits within the jupiter database. similarly, geotechnical boreholes, which are drilled prior to the construction of buildings and infrastructure, have not been systematically used to identify interglacial deposits. in denmark, boreholes that are not considered to affect groundwater resources, including boreholes for geotechnical investigation, raw material prospecting or scientific research, are typically excluded from mandatory reporting to jupiter. therefore, geotechnical data are commonly stored only by industry actors in private archives, and the findings of interglacial deposits remain publicly unreported. some sectors of the industry have recently made archived data more available, such as the geoatlas live solution by geo, where simplified borehole data are displayed and shared between stakeholders and subscribers. this study aims to assess the capacity of geotechnical borehole data to identify accessible interglacial deposits that may hold traces of early hominin presence in denmark. our approach concentrates on compiling data on well-documented interglacial lacustrine deposits from private geotechnical databases. the specific aims are, firstly, to locate glacially undisturbed mis 5 lacustrine deposits with a focus on southern and southwestern jylland, and secondly, to demonstrate the potential of large geotechnical data sets. combined, this draws the contours of the potential of geotechnical data in both earth and palaeoenvironmental research, provided that general issues of open access and open data can be resolved. 2 materials and methods 2.1 geotechnical borehole data geotechnical investigations in denmark follow a common standard formulated by the danish geotechnical society (felthåndbogen 1999). this means that geotechnical boreholes are commonly drilled with an 8″ cased auger, with samples taken every 0.5 m; however, this can differ depending on ground conditions, purpose and companies. since the late 1990s, borehole data have predominantly been managed within geogis, a database software developed in denmark and designed for handling geological, geotechnical and water-related data and tasks. the use of the geogis software has facilitated the digitalisation and accumulation of data in large internal databases over time. borehole data are arranged in ‘projects’, with referred ‘points’, with referred ‘samples’, that again have ascribed attributes, for example, sediment description, depth, moisture content, etc. the sample description follows a common danish standard (larsen et al. 1995), where the sediment is classified according to its physical appearance by main component(s), grain size, sorting or grading, secondary components, structures, colour and calcareousness. the depositional environment of each sample is interpreted and assigned an age with a code, for example ‘fe ig’ stands for ‘freshwater (fe) environment and interglacial (ig) age’. the interpretation is based on the geology and the overall stratigraphy. the geotechnical https://doi.org/10.34194/geusb.v57.8372 http://www.geusbulletin.org/ nielsen et al. 2024: geus bulletin 57. 8372. https://doi.org/10.34194/geusb.v57.8372 3 of 10 geusbulletin.org description does not distinguish palustrine sediments, and these would commonly be referred to a freshwater environment. in cases of doubt or redeposition, multiple interpretations can be given, for example ‘fe pg/ig’ (pg: post-glacial). commonly, the water content of cohesive sediment samples is measured along with the in situ ground strength. with a focus on south and south-west jylland, six major geotechnical companies were contacted: geosyd, franck geo-og miljøteknik, sweco, rambøll, niras and geo. all companies granted access to their databases. the six geogis databases were accessed from september 2022 to may 2023. ownership and rights regarding the data differ from company to company and are generally informal, with many companies practicing shared ownership with their clients. we did not experience any refusal to share data, but concerns from the six companies were raised, leading to limitations in the display of the data. to protect company clients from third-party intervention, information that could identify clients has been anonymised in this paper. access to the original data can only be granted by the data owners. relevant project references are provided in supplementary file s2 or can be obtained from the corresponding author on request. 2.2 compilation of geotechnical borehole background data to evaluate the overall coverage and test geotechnical borehole data against known interglacial sites, all available borehole data from five of the six databases were compiled with an sql search at the ‘point’ level. a total of 422 072 points were compiled and evaluated. the raw quantities of the databases are shown in table 1. the point data were then compiled into one data set at the ‘project’ level, attributed with the number of boreholes. there are some inherited differences between the databases where some attributes may not have been entered in every database and coordinates have sometimes been handled separately. all available coordinates were transformed to utm zone 32n (epsg: 25832) and plotted in qgis. in this process, several common errors were identified in the databases, such as erroneous projection, transverse coordinates, decimal and typographical errors. obvious errors were, to a large degree, corrected manually. where project coordinates were not available, a coordinate was calculated as a mean of point coordinates. again, in databases where project titles often resembled addresses, and point coordinates were not available, the project title was georeferenced using the qgis plugin, danish address tool by ©septima. one database contained imports from the publicly available data set jupiter and from municipalities, and these were excluded. projects identified as being located outside denmark were excluded from the data set. due to limitations in manual validation of all project positions, we acknowledge the potential introduction of errors at various stages. associating all boreholes within a project to a single coordinate can lead to errors, particularly in cases involving multiple locations or referencing a linear shape. during the analysis of various databases, it became evident that certain projects were shared among multiple companies or databases. however, the scale of this issue was assessed to be negligible, and systematic filtering of duplicate background data was not implemented. all projects were plotted in qgis, enabling a geographical search, and a density map of all boreholes was generated using a 10-km search radius for each cell and weighing the projects by number of points (boreholes), see fig. 1. the island of bornholm was poorly represented in the data and has been excluded from further analyses, together with any offshore boreholes. 2.3 compilation of interglacial and interstadial samples in the six geotechnical databases, a separate sql search was performed to identify all samples with an attributed interglacial or interstadial (igis) age. the data were exported to spreadsheets, with a total of more than 11 000 igis samples from more than 1800 boreholes. imported raw quantities are shown in table 1. table 1 overview of database sizes and the inherent interglacial or interstadial (igis) points (boreholes) and samples. database date of access projects points samples (dd.mm.yyyy) all igis all igis igis geosyd 16 september 2022 46 838 302 161 426 600 1990 niras 11 november 2022 1750 21 24 598 66 533 sweco 20 december 2022 7919 75 69 023 170 891 franck 10 january 2023 10 962 45 59 806 147 1521 rambøll 31 january 2023 5386 113 107 219 361 2498 geoa 10 may 2023 unknown 144 >147 000 514 4345 sum >72 855 700 >569 072 1858 11 778 a access has only been granted to igis boreholes. https://doi.org/10.34194/geusb.v57.8372 http://www.geusbulletin.org/ http://dd.mm nielsen et al. 2024: geus bulletin 57. 8372. https://doi.org/10.34194/geusb.v57.8372 4 of 10 geusbulletin.org all igis samples were compiled into boreholes (n  =  1858; 0.33% of total points), georeferenced and plotted in qgis (see section ‘compilation of geotechnical borehole background data’). all points underwent manual validation; however, eight boreholes could not be georeferenced and were excluded. to identify lacustrine deposits of significant size and reliable data among igis boreholes (totalling n  =  1850), a filtering process consisting of four sequential steps was conducted within a spreadsheet. the filtering ascribed the boreholes to five distinct categories listed below. a detailed protocol for the filtration process is accessible in the supplementary file s1. 1. category i: highly unlikely or errors (n = 63): this category encompasses boreholes where igis samples could be classified as first (a) topsoil (n = 35) or then (b) duplicates (n = 28). 2. category ii: inconclusive or sparse data (n = 912): this category includes boreholes where (a) the majority of samples had a secondary age not being interglacial, interstadial or postglacial (n  =  437). additionally, it comprises boreholes with (b) five or fewer samples within a 250 m range (n  =  252) or (c) two or fewer consecutive samples (n = 223). 3. category iii: few samples or an inorganic context (n = 587): this category comprises (a) boreholes with less than 15 samples within a 250 m range (n = 169) and (b) boreholes with few or no indications of organic content (n = 418). 4. category iv: inaccessible or marine deposits (n = 127): this category includes boreholes where (a) igis samples are buried more than 10 m beneath the surface, making them difficult to access (n = 71). additionally, it covers boreholes where (b) the majority of samples have been interpreted as marine (n = 56). 5. category v: boreholes indicating significantly large deposits of organic-rich freshwater sediment, fairly sure to be of interglacial or interstadial age (n = 161). a graphical presentation of the filtration and classification for this study is shown in fig. 2. all 161 category v igis points were combined into 39 igis occurrences together with neighbouring igis points. a map of the occurrences is shown in fig. 3. the localities were crosschecked with localities known from literature and nearby jupiter boreholes. all sites were attributed an occurrence number, occurrence name and (1) source information: number of category v and ii–iv boreholes; (2) spatial information: utm32n coordinate, maximal and minimal elevation and maximal and minimal depth; (3) short comment with characteristics, observations and related geomorphology; (4) references to relevant literature, relevant nearby jupiter boreholes and references to the geotechnical database and adjoining project number. the attributes for all 39 occurrences are shown in supplementary file s2. 2.4 jupiter data for comparison, borehole data from the jupiter database have been assessed. a search of the boreholes with digitised lithology, june 2024, revealed 1820 boreholes containing interglacial deposits, 514 of which contained interglacial freshwater sediments of likely terrestrial origin. when filtered similarly to the geogis data, with a depth less than 10 m, total thickness of interglacial layers greater than 0.5 m and sediment to be of gyttja or peat, only 58 boreholes remain. due to the huge difference in recorded parameters between the geotechnical and jupiter databases, however, boreholes from jupiter have not been included in this study and have not been filtered for the concentration criteria (more than 15 igis samples within a 250 m range). 3 results and discussion 3.1 marine deposits three well-known marine interglacial deposits (knudsen 1995) appear frequently in the databases. these are (1) fig. 1 density map illustrating the distribution of geotechnical projects, weighted by the number of points (boreholes), showing borehole density within a 10-km radius. the data were sourced from five of the six databases used in this study, encompassing more than 450 000 boreholes from more than 72 000 projects. the map indicates a significant bias in the data set towards urbanised areas, particularly in the greater copenhagen area and east jylland. the main stationary line during the last glacial maximum is depicted by the dashed blue line, while the remnant pleistocene surface of the danish hill-islands is represented with a grey mesh, modified from jakobsen (2022). 12°e 15 000 7500 0 jylland vendsyssel borehole density 10°e8°e 57°n 55°n 0 50 100 km sjælland fyn sw baltic sea n or th s ea copenhagen https://doi.org/10.34194/geusb.v57.8372 http://www.geusbulletin.org/ nielsen et al. 2024: geus bulletin 57. 8372. https://doi.org/10.34194/geusb.v57.8372 5 of 10 geusbulletin.org the eemian cyprina clay (kristensen & knudsen 2008), familiar to the south-western baltic sea, and often found incorporated in weichselian tills in the aabenraa, haderslev and south fyn area. in some company records, this organic-rich marine clay is referred to as “eemler”, eemian clay. (2) the marine eemian and early weichselian skærumhede series known from northern jylland (pedersen 2005), also referred to as the “older yoldia clay”. this deep marine deposit is found in the vendsyssel area in outcrops and deep boreholes. (3) marine deposits are also known from older interglacial stages, especially in south-western jylland, where the likely elsterian “oldest yoldia clay” and subsequent holsteinian marine deposits are found in many boreholes around the city of esbjerg and have been described from former clay pits there. 3.2 significant interglacial and interstadial occurrences all 39 identified igis occurrences are shown on the map in fig. 3 and listed in supplementary file s2. in some cases, additional data were obtained from the companies in the form of full geotechnical or excavation reports although this was not always possible. of the 39 occurrences identified in this study, two were already known and described in the literature: the brørup stationsby deposit (occurrence (occ.) 25; hartz 1909) and the kollund lake deposit at lind (occ. 15; kronborg & odgaard 2004; odgaard & møbjerg 2009). five occurrences are in proximity to, or can be correlated with, known deposits, yet the added observations are substantial (occ. 18, 20, 21, 24 and 35). a total of 32 of the identified occurrences appear to be new findings of igis deposits. nine of the 39 occurrences are selected and described below to illustrate and discuss the potential of the identified igis deposits. 3.2.1 hvorup (occ. 1) located just north of nørresundby, north jylland, this is the northernmost site with an interglacial freshwater deposit found in this study and known from the literature. the deposit is located under the gently sloping surface of late glacial younger yoldia sand (berthelsen 1987) on the eastern flank of hvorup hill. the deposit is buried 3.5–4.5 m below the surface (b.s.) and is more than 8.5 m thick. it is recorded in six boreholes within a 30 × 30 m quadrant and may correlate with two jupiter boreholes within 130 m. samples are described mainly as peaty gyttja, with more clayey horizons towards the base. no similar freshwater interglacial deposits have been reported from this area, and the age of the deposit remains unresolved. 3.2.2 ans (occ. 2) the ans igis deposit is located 10 km north of silkeborg, in central jylland. the interglacial deposit was encountered during the development of a new residential area. identified by two separate companies in three different projects, the deposit consists mainly of gyttja with diatomite, up to 6 m thick. one investigation recorded high calcareous content, which is generally rare for interglacial deposits. the deposit is situated within an area of 100 × 100 m on a gentle slope within the extended south-eastern course of a buried valley, identified as area aar44 in sandersen and jørgensen (2016). the age of the deposit is unresolved, but if it is found to be an infill of the buried valley, it would likely be pre-eemian. fig. 2 flow diagram illustrating the filtration of boreholes into categories i–v, based on sediment descriptions, quantity, interconnectivity, accessibility and more. a total of 161 boreholes were identified as containing significantly large deposits of organic-rich freshwater sediment with interglacial or interstadial age. these were grouped into 39 occurrences, incorporating neighbouring boreholes from the categories ii–iv. diagram created using sankeymatic. n = 1 85 0 i (n = 63) ii (n = 912) iii (n = 587) iv (n = 127) v (n = 161) https://doi.org/10.34194/geusb.v57.8372 http://www.geusbulletin.org/ nielsen et al. 2024: geus bulletin 57. 8372. https://doi.org/10.34194/geusb.v57.8372 6 of 10 geusbulletin.org germany silkeborg herning esbjerg lgls occurrence iv borehole iii borehole ii borehole sweden germany copenhagen roskilde odense aabenraa haderslev kolding fredericia vejle horsens aarhus randers 1 2 3 4 5678 9 10 1213 14 15 1617 18 19 20 21 2223 24 2526 27 28 29 30 31 32 33 34 35 36 37 38 12°e11°e10°e9°e8°e 57°n 56°n 55°n aalborg 39 0 25 50 75 100 km fig. 3 the distribution of boreholes categorised as ii, iii and iv is shown in greyscale. identified interglacial and interstadial (igis) occurrences are highlighted in red, and locations discussed in the text are labelled in bold. details for all 39 occurrences are provided in supplementary file s2. the labelled occurrences are: 1: hvorup, 2: ans, 3: engskovgård, 4: birkemose, 5: engesvang, 6: holing, 7: fuglsang, 8: gødstrup, 9: snejbjerg, 10: vildbjerg i, 11: vildbjerg ii, 12: møltrup, 13: videbæk, 14: lem i, 15: kollund sø, 16: elmholt, 17: flø, 18: risbanke ii, 19: dagnæs, 20: mølholm valley, 21: karensdal, 22: kolding v, 23: gesten, 24: brørup n, 25: brørup stationsby, 26: holsted stationsby, 27: bramming i, 28: kjersing, 29: rørkjær, 30: gredstedbro, 31: møgelmose, rødding, 32: toftlund i, 33: bådsbøl, 34: tønder n, 35: tønder, 36: horsbyg, 37: havnbjerg, 38: nordborg, 39: sdr. vedby. https://doi.org/10.34194/geusb.v57.8372 http://www.geusbulletin.org/ nielsen et al. 2024: geus bulletin 57. 8372. https://doi.org/10.34194/geusb.v57.8372 7 of 10 geusbulletin.org 3.2.3 vildbjerg i (occ. 10) a concentration of 12 boreholes from four different projects is located south-west of vildbjerg, central jylland. here, interglacial freshwater gyttja and peat are found within an area of approximately 200 × 400 m situated on a gentle valley slope. the deposit is found at 3.5–21.5 m b.s. and is up to 16.5 m thick. it appears homogeneous with a moisture content typically ranging from 50 to 100% and with a few peat layers reaching 200%. the deposit is not recorded in nearby jupiter boreholes. the age is unknown, but based on a geomorphological assessment, a pre-eemian age is suggested. 3.2.4 møltrup (occ. 12) initial observations from geotechnical boreholes and reports indicated highly irregular deposits of thin peat layers, thick layers of gyttja, clay and silt, among sorted sand with variable organic content. part of the site development project was apparently abandoned in areas with the most organic soils. however, a subsequent excavation unveiled significant deformation with organic-rich soil a and b horizons, exhibiting a sub-vertical orientation in the profile. the combined findings strongly suggest the presence of a glaciotectonically disturbed soil surface with a thick podzol, with a combined thickness of o-a-e-bh-bs-horizons exceeding 1.5 m, developed from glacially derived parent material. no published examples of pleistocene palaeosols from south-west scandinavia (sjørring & frederiksen 1980; stephan 2014; nielsen et  al. 2019; houmark-nielsen 2021), nor from  pliocene or miocene deposits in denmark (rasmussen 2017), contain such ‘giant’ podzols. in this regard, the møltrup site represents an anomalous and complex interglacial deposit that merits further investigation. 3.2.5 elmholt (occ. 16) situated on the skovbjerg hill-island, this interglacial deposit was encountered prior to the construction of a highway south of herning, central jylland. the deposit has a thickness of more than 12 m and extends for 400 m in a north–south direction, while its east–west extent remains unknown. it consists mainly of non-calcareous diatomic peat and gyttja with water content of 80–110% and is overlain by clay with a high moisture content, approximately 35%. the age is undetermined, but similar deposits in the area have been correlated to mis 11 (odgaard et al. 2016) and a pre-eemian age aligns well with the geomorphology at the site. 3.2.6 risbanke ii (occ. 18) a single deep borehole penetrated a thick series of apparently undisturbed glacial and interglacial lake sediments. the bottom of the series starts at 17 m b.s. with slightly organic and slightly calcareous, then non-calcareous clay, shifting to a varved, very calcareous silt at 13 m b.s., followed by gyttja, which becomes non-calcareous. at 8.5 m b.s., the deposit consists of laminated clay, which is followed by gyttja at 5 m b.s. and then by clay. the igis deposit is within 500 m of a known interglacial site, risbanke (i), ascribed to the eemian interglacial (milthers 1939), but the deposits cannot be directly correlated. 3.2.7 mølholm valley (occ. 20) within and in front of mølholm valley, a side valley to vejle fjord valley, east jylland, several boreholes have encountered a deposit of freshwater sediments. the deposit consists of homogeneous diatom gyttja, which is more than 31 m thick. it is situated within the buried valley of randsfjord (rib17; sandersen & jørgensen 2016), stretching nearly 25 km from vejle south-east to fredericia. the interglacial site of vejlby within this valley is ascribed to the holsteinian interglacial (andersen 1965). the deposit at mølholm valley might not be novel, but it represents the thickest section found so far and could provide a high-resolution archive of the holsteinian interglacial. 3.2.8 kolding v (occ. 22) the deposit is located within the ‘kolding v’ (kolding west) highway intersection, but the exact location has not been georeferenced, as the data originate from an investigation carried out in 1968. the data have since been digitised by the company geo. five boreholes penetrated interglacial gyttja, which appeared near the surface or buried 5 m below, with a thickness of up to 7.5 m. the deposit consists of clay at the base, overlain by gyttja. the upper part consists of gyttja, silt and sand layers. the age of the deposit is unknown. 3.2.9 gredstedbro (occ. 30) the igis deposit at gredstedbro is documented by single investigation and situated on a high-lying terrace on the western side of the holsted hill-island, to the northeast of gredstedbro. here, five boreholes revealed an interglacial deposit primarily composed of homogeneous sand and interbedded freshwater gyttja, at depths ranging from 1.3 to 4.9 m b.s. the gyttja, likely lacustrine in origin, is described as having coal-like sections and exhibits a high moisture content of approximately 100–150%. this gyttja layer overlays well-sorted sand with plant remains and is covered by well-sorted medium-fine sand. the age of the deposit remains undetermined. https://doi.org/10.34194/geusb.v57.8372 http://www.geusbulletin.org/ nielsen et al. 2024: geus bulletin 57. 8372. https://doi.org/10.34194/geusb.v57.8372 8 of 10 geusbulletin.org 3.3 poor coverage of pleistocene basins the borehole distribution map, fig. 1, reveals a clear bias towards developed areas around denmark’s larger cities. in contrast, west jylland, west of the main stationary line of the last glacial maximum, shows relatively sparse coverage of geotechnical boreholes. this likely reflects the limited recent development following low population density in the region. the low degree of urban development especially applies to present-day depressions, which are left undeveloped where possible. a notable example is the eemian palaeolake at the herning clay pit, first described by jessen and milthers (1928). situated in a weak depression measuring approximately 250 × 800 m, this palaeolake is now located within a densely populated area. the area has been maintained as a recreational area, with football fields and only a few roads and a railroad line crossing it. this tendency to avoid developing basins on the remnant pleistocene surface in urban areas appears to be common and reflects an avoidance of wetlands and additional cost associated with constructing deeper foundation. the site of herning clay pit also highlights a general reluctance to ascribe deposits to igis ages when organic layers are not buried by subsequent glaciogenic deposits. in 2018, a geotechnical investigation in the south-eastern corner of the herning clay pit (madsen 2018a) initially misinterpreted the dark organic-rich layers as fill, glaciogenic or of other origins. only upon deeper reinvestigation were these deposits correctly identified as interglacial (madsen 2018b). 3.4 limitations of geotechnical data in filtering relevant boreholes (fig. 2), we set a high threshold for boreholes, leading to the exclusion of some promising boreholes with limited neighbouring igis boreholes. this threshold, especially the criterion of more than 15 samples within a 250 m range, presents issues for large infrastructure projects where borehole density is low and often aligned in transects rather than clusters. to address this, we could prioritise the thickness of igis layers in individual boreholes over the number of neighbouring samples. however, for this initial study, we considered the spatial extent obtained from multiple boreholes more important for characterising the deposits. another potential improvement involves expanding the data set beyond the six geotechnical databases accessed for this study, particularly focusing on west jylland. including other geogis databases or jupiter data or addressing other research questions could enhance our understanding and utilisation of these data. 3.5 catching false negatives the accuracy of sediment classification and geological interpretation by technicians heavily influences our methodology. as illustrated by the herning clay pit example (section ‘poor coverage of pleistocene basins’), we suspect that many organic-rich interglacial deposits, especially those not overtly covered by glaciogenic sediments, are mistakenly attributed a postglacial age. this misattribution often arises from reliance on stratigraphy, where interglacial age assignment depends on the presence of overlying glaciogenic deposits. this issue is particularly pronounced when igis deposits are buried beneath holocene lacustrine deposits, where there is only a subtle transition between the postglacial and interglacial layers. during this study, we considered using moisture content as a proxy for age, based on the idea that sediment compacts over time due to burial, permafrost processes and organic material decomposition. however, the heterogeneous nature of the parent material, coupled with other variables such as overlay pressure, drainage patterns and decomposition, renders moisture content an unreliable proxy for age. 3.6 potential for preservation in regions the preservation of organic material is essential to establish a robust archive to reconstruct palaeoenvironments. organic deposits typically indicate an anoxic environment, which protects them from microbial decay. although comprehensive evaluations of organic degradation are challenging, insights can sometimes be gleaned from sample descriptions, especially when macrofossils are identified. for peat, the dgf standard (larsen et al. 1995) outlines four levels of degradation, yet these are not consistently applied in sediment descriptions. another crucial factor affecting preservation is the ph level, which directly impacts bone preservation – a key consideration in fauna studies and archaeological investigations involving cutmarks and bone modifications. high ph levels are also known to support ancient dna (adna) preservation (giguet-covex et al. 2019). while geotechnical sediment descriptions do not include ph measurements, hcl testing can determine the presence of calcareous minerals. likewise, the presence of shells can provide clues regarding a neutral or alkaline environment. among the 39 occurrences found in the survey, 10 have recorded calcareousness, with just two sites – ans (2) and risbanke ii (18) – documenting partially calcareous organic deposits. in some cases, subsequent glaciations have influenced interglacial deposits to varying degrees. https://doi.org/10.34194/geusb.v57.8372 http://www.geusbulletin.org/ nielsen et al. 2024: geus bulletin 57. 8372. https://doi.org/10.34194/geusb.v57.8372 9 of 10 geusbulletin.org glaciotectonic disturbances can significantly disrupt these deposits, compromising their structural integrity. detecting glaciotectonic structures through geotechnical boreholes alone is challenging, as exemplified by the møltrup site (occ. 12), see section 3.2.4. 3.7 mis 5 or older deposit most interglacial sites in denmark have been referred to either the eemian (mis 5e) or the holsteinian (mis 11c) stages, but other interglacial periods are also represented. although dating of the identified igis deposits is beyond the scope of this study, we evaluated the 39 occurrences based on the present geomorphology, historical topographic maps, overall descriptions and related literature. from this evaluation, we believe that it is unlikely any deposits relate to the late glacial interstadials bølling-allerød. our criteria – more than two consecutive samples and a minimum of 15 samples within close proximity – appear effective in this regard. none of the 39 localities could be directly linked to depressions in the pleistocene surface, unlike several known examples of mis 5 deposits, such as the brørup bogs (jessen et al. 1918) or the herning clay pit (jessen & milthers 1928). this could be due to terrain alterations following construction. however, several deposits present at current terrain levels and located at the pleistocene surface are likely of eemian age, including occurrence 5–9, 13–14, 17 and 23–32. a few occurrences correlate with known deposits older than the eemian, this being occurrences 15 and 20–21. some deposits appear in thicknesses and distributions hitherto unprecedented in an eemian context, suggesting a pre-eemian age, including occurrence 10–11, 16 and 18. due to limited information, we refrain from ascribing any age estimates to the remaining localities. as for all localities, reliable dates should be obtained with suitable methods if further studied. 4 conclusion this study identified 39 new interglacial sites, many of interest to the scientific community. the method of querying private databases has proven particularly effective in locating igis deposits, with evident glacial overprints. these sites require further study to fully explore their potential and to date these deposits properly. our data set included many marine deposits, which were not discussed here, but could be crucial climate archives and lead to a better understanding of historical sea-level fluctuations. supplementing this data set with additional geogis databases is feasible, and new data are continuously produced. thus, this approach offers significant promise for identifying unknown interglacial sites in the future. however, this approach is limited in identifying sites suitable for exploring early hominin evidence. determining highly attractive campsites requires an understanding of the interconnectivity to the broader palaeolandscape. the identified sites generally lack geomorphological context. to address biases in geotechnical borehole data away from pleistocene depressions (see section ‘poor coverage of pleistocene basins’) and avoid false postglacial attributions (see section ‘catching false negatives’), new strategies are necessary to locate undisturbed eemian deposits on a large scale. this study demonstrates, for the first time in denmark, how privately owned geotechnical databases can be used for scientific inquiries when made accessible. many more applications are conceivable. open access to archived and ongoing geotechnical data would benefit various fields, including archaeology, quaternary geology, groundwater mapping and environmental risk and geohazard assessment. a model that respects owner rights could include only sediment descriptions, groundwater tables and layer boundaries, while omitting quantitative data on the soil geotechnical properties. acknowledgements we extend our gratitude to the companies geosyd (haderslev), franck geoog miljøteknik (horsens), sweco (kolding), niras (allerød), ramboll (aarhus) and geo (aarhus), for granting access to their geogis-databases and for their interest and collaboration in this project. special thanks to ramboll for providing a licence and technical support for the geogis2020 software and to the fællesbo social housing organisation for sharing the geotechnical reports from teglparken, herning. we appreciate the critical evaluations and recommendations from three reviewers, which have improved this manuscript. the clioarch team is also thanked for their valuable discussions and comments on the figures and language of this paper. additional information funding statement the presented work has been founded by the independent research fund denmark as a part of the neanderedge project (case number 9062-00027b). author contributions emsn: conceptualisation, resources, data curation, formal analysis, investigation, methodology, validation, visualisation, writing – original draft, writing – review and editing; smk and tkn: conceptualisation, supervision, writing – review and editing; tkn: funding acquisition, project administration. competing interests the authors declare no competing interests additional files three supplementary files are available at https://doi.org/10.22008/ fk2/48hra5. these include: (1) a readme file. (2) supplementary file s1: protocol of data processing and filtration of igis boreholes. 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https://doi.org/10.1038/s41598-024-57490-4 applying geotechnical borehole databases in the search for interglacial deposits in denmark 1 introduction 2 materials and methods 2.1 geotechnical borehole data 2.2 compilation of geotechnical borehole background data 2.3 compilation of interglacial and interstadial samples 2.4 jupiter data 3 results and discussion 3.1 marine deposits 3.2 significant interglacial and interstadial occurrences 3.2.1 hvorup (occ. 1) 3.2.2 ans (occ. 2) 3.2.3 vildbjerg i (occ. 10) 3.2.4 møltrup (occ. 12) 3.2.5 elmholt (occ. 16) 3.2.6 risbanke ii (occ. 18) 3.2.7 mølholm valley (occ. 20) 3.2.8 kolding v (occ. 22) 3.2.9 gredstedbro (occ. 30) 3.3 poor coverage of pleistocene basins 3.4 limitations of geotechnical data 3.5 catching false negatives 3.6 potential for preservation in regions 3.7 mis 5 or older deposit 4 conclusion acknowledgements additional information funding statement author contributions competing interests additional files references figures fig. 1 density map illustrating the distribution of geotechnical projects, weighted by the number fig. 2 flow diagram illustrating the filtration of boreholes into categories i-v, based on sediment fig. 3 the distribution of boreholes categorised as ii, iii and iv is shown in greyscale. identified table table 1 overview of database sizes and interglacial or interstadial (igis) ages. points are the num geological survey of denmark and greenland bulletin 11, 115-123 115 new hornblende and muscovite 40ar/39ar cooling ages in the central rinkian fold belt, west greenland ann-sofie sidgren, laurence page and adam a. garde the palaeoproterozoic rinkian fold belt in west greenland consists of reworked archaean basement, mainly orthogneiss, and the unconformably overlying palaeoproterozoic karrat group. both parts were intensely deformed and metamorphosed at around 1.87 ga, at which time the crustal anatectic prøven igneous complex was emplaced into the northern part of the belt. seven new hornblende and muscovite 40ar/39ar cooling ages are presented from the central–northern parts of the rinkian fold belt. four 40ar/39ar hornblende ages ranging from 1795 ± 3 to 1782 ± 3 ma were obtained from amphibolite and hornblendite enclaves in the archaean orthogneiss, and two from relict dyke fragments in the latter that may be of palaeoproterozoic age. three 40ar/39ar muscovite ages of 1681 ± 6 ma, 1686 ± 3 ma and 1676 ± 3 ma were obtained from samples of karrat group metagreywacke, andalusite schist and metasiltstone. the new 40ar/39ar ages, from hornblende and muscovite respectively, are very uniform and probably unrelated to local metamorphic grade and structural history, and are interpreted as regional late orogenic cooling ages. the new hornblende ages are significantly older than those previously obtained from the central and northern parts of the adjacent nagssugtoqidian orogen to the south, and point to different uplift histories, which may suggest that the orogeny was not synchronous in the two regions. keywords: ar-ar, geochronology, rinkian, palaeoproterozoic, west greenland ____________________________________________________________________________________________________________________________________________________________________ a.-s.s. & l.p., department of geolog y, geobiosphere science center, lund university, sölvegatan 12, s-223 62 lund, sweden. e-mail: ann-sofie.sidgren@bd.lst.se a.a.g., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. this paper presents seven new hornblende and muscovite 40ar/39ar cooling ages from the central part of the palaeoproterozoic rinkian fold belt in west greenland. the new data set provides insight into the cooling history of the rinkian fold belt and can also be used to address its temporal relationship with the adjacent nagssugtoqidian orogen to the south, from which other 40ar/39ar cooling ages have previously been published. most of central and northern west greenland consists of archaean continental crust which was intensively reworked during the palaeoproterozoic. this reworking was first recognised in central west greenland between 66° and 69°n by ramberg (1949), who established the nagssugtoqidian mobile belt (now called the nagssugtoqidian orogen) in this area. escher & pulvertaft (1976) subsequently proposed that a separate palaeoproterozoic mobile belt, the rinkian fold belt, existed in central and northern west greenland between 69° and 75°n (see inset map of fig. 1). they noted that the latter region was dominated by an overall flat-lying tectonic foliation with superimposed large domes, and considered that these structures were of a different nature from the generally steep foliations and tight folds that had previously been identified in the nagssugtoqidian belt. in contrast to the collisional structures recognised within the nagssugtoqidian orogen, it was thought that the rinkian deformation had taken place without significant crustal shortening. furthermore, whereas the collisional nagssugtoqidian © geus, 2006. geological survey of denmark and greenland bulletin 11, 115–123. available at: www.geus.dk/publications/bull 116 orogen was originally believed only to comprise archaean supracrustal and infracrustal rocks, the rinkian fold belt contains a widespread, metamorphosed and deformed cover sequence, the c. 2 ga old karrat group, which was unconformably deposited on the archaean basement gneisses (garde & pulvertaft 1976; henderson & pulvertaft 1987; kalsbeek et al. 1998). the lowest parts of the karrat group consist of quartzite, marble and minor amphibolite (the qeqertarsuaq and marmorilik formations), which are overlain by a very uniform sequence of metagreywacke, the nukavsak formation, which is several kilometres thick and occurs throughout most of the rinkian belt (fig. 1; henderson & pulvertaft 1987). the geochemistry of the karrat group and studies of its detrital zircons indicate that the karrat group was derived from a mixed source including palaeoproterozoic magmatic arc 483667 hbl 1782 ± 2 ma 483696 hbl 1785 ± 3 ma 483708 hbl 1784 ± 3 ma 483657 hbl 1795 ± 3 ma 483671 mu 1676 ± 3 ma 483654 mu 1686 ± 3 ma 483653 mu 1681 ± 6 ma 500 km rinkian nagssugtoqidian disko bugt 72°n 51°w 55°w 71°n upernavik kangilleq 50 km ukkusissat fjord qeqertarsuaq qinngusaaq inland ice surficial deposits quaternary basalt paleocene sandstone karrat group palaeoproterozoic prøven igneous complex archaean, variably reworked orthogneiss and minor supracrustal rocks ar-ar sample hbl: hornblende mu: muscovite kangerlussuaq uummannaq appat svartenhuk nuussuaq fig. 1. map of the central rinkian fold belt showing the locations of samples collected for 40ar/39ar age determination. the index map shows the position of fig. 1 and the approximate extent of nagssugtoqidian and rinkian reworking in west greenland. modified from garde et al. (2004). 117 rocks and archaean basement rocks, and that it was deposited at around 2.0–1.9 ga ago (kalsbeek et al. 1998; thrane et al. 2003). the rinkian fold belt also incorporates the prøven igneous complex, a very large plutonic complex of granitic and microdioritic crustal melts that were emplaced under granulite facies conditions into the middle to upper crust in the upernavik region of the rinkian belt and is also found as the cumberland batholith on adjacent baffin island, canada. the pluton has previously yielded a rbsr whole rock isochron age of 1860 ± 25 ma with a high initial 87sr/86sr ratio (kalsbeek 1981) and has recently also been studied by thrane et al. (2005). the latter produced a more precise zircon u-pb ion probe age of 1869 ± 9 ma and obtained negative εnd values from the pluton (calculated at 1870 ma) ranging between –5.2 and –4.3. in agreement with the previous rb/sr data this shows that the plutonic complex contains a large archaean continental crustal component. it is therefore questionable whether the prøven igneous complex – cumberland batholith is subduction-related as has been proposed by canadian workers. thrane et al. (2005) suggest it represents a crustal melt, induced by upwelling hot asthenospheric mantle. pulvertaft (1986), henderson & pulvertaft (1987), grocott & pulvertaft (1990) and garde & steenfelt (1999b) have described the structural evolution in various parts of the rinkian fold belt, and recognised largescale thrusts in its southern part. following new field work in 2002–2003, the structural evolution in the uummannaq region is at present regarded as consisting of four main phases (briefly outlined by garde et al. 2003, 2004). deformation began with tight folding and possibly thrusting (d1), which developed prior to cleavage formation. this was followed by neto e-directed thrusting and ductile tectonic transport (d2) accompanied by formation of a penetrative schistosity, and then by nwto w-directed tectonic transport (d3) and intensification of the pre-existing schistosity. lastly, very large, upright to overturned, dome-shaped anticlines and tight synclinal cusps were developed during continued shortening of the now strongly tectonically layered crust; these large structures are only locally accompanied by a new tectonic fabric. the prøven igneous complex was emplaced at a late stage of the main fabric-forming events and gave rise to a wide metamorphic aureole that was overprinted on rocks that were already regionally metamorphosed at high grade. 1900 1700 0 4020 60 80 100 1750 1850 1800 1795 ± 3 ma integrated age = 1797 ± 3 ma hornblende483657 1800 1600 1650 1750 1700 1681 ± 6 ma* integrated age = 1703 ± 3 ma muscovite483653 1600 1650 1750 1700 1686 ± 3 ma* integrated age = 1716 ± 2 ma muscovite483654 1600 1650 1750 1700 1676 ± 3 ma integrated age = 1684 ± 3 ma muscovite483671 1700 0 4020 60 80 100 1750 1850 1800 1782 ± 2 ma integrated age = 1801 ± 2 ma hornblende483667 1900 1700 1750 1850 1800 1785 ± 3 ma integrated age = 1864 ± 2 ma hornblende483696 1700 0 4020 60 80 100 1750 1850 1800 1784 ± 3 ma integrated age = 1798 ± 3 ma hornblende483708 cumulative % 39ar released a pp ar en t ag e (m a) ukkusissat area north-eastern uummannaq area south-eastern uummannaq area fig. 2. 40ar/39ar plateau age spectra from the rinkian fold belt. ages with an asterisk (*): 40ar/39ar plateau age representing less than 50% of total 39ar release. ages without an asterisk: 40ar/39ar plateau age. 118 483653 muscovite (j = 0.01071 ± 0.00001): c 1.4 0.00001 0.001 0.00011 0 147.485 0.093 10.9 10.9 100 1709.6 4.4 d 1.4 0.1461 0.004 0.00008 26.5 147.544 0.306 35.6 46.4 100 1710 1.7 •e 1.5 1.9453 0.021 0.00097 27.6 143.552 0.011 1.2 47.7 99.9 1679.9 13.4 •f 1.6 2.875 0.004 0.00041 96.1 143.488 0.009 1 48.7 100 1679.4 17.2 •g 1.9 0.4853 0.003 0.00010 68.7 143.739 0.187 21.7 70.3 100 1681.3 2.7 h 4 0.3402 0.004 0.00019 24.8 147.664 0.256 29.7 100 100 1710.9 1.6 integrated (total fusion) age: 1703 3 (•) plateau age: 23.9 1681 6 483654 muscovite (j = 0.01071 ± 0.00001): a 1.4 0.005 0.003 0.00106 0.1 150.444 0.767 26.9 26.9 99.8 1731.6 1.0 b 1.5 0.006 0.005 0.00033 0.2 149.670 0.692 24.3 51.2 99.9 1725.9 1.2 •c 1.6 0.0499 0.009 0.00063 1.1 144.382 0.265 9.3 60.5 99.9 1686.2 1.5 •d 1.7 0.9006 0.020 0.00077 16.1 143.337 0.075 2.6 63.1 99.9 1678.2 3.3 •e 1.8 0.1521 0.007 0.00147 1.4 144.323 0.153 5.4 68.5 99.7 1685.7 2.2 •f 1.9 0.357 0.009 0.00015 32.9 144.706 0.251 8.8 77.3 100 1688.6 3.1 g 2.3 0.0741 0.003 0.00003 35.7 148.939 0.648 22.7 100 100 1720.4 2.1 integrated (total fusion) age: 1716 2 (•) plateau age: 26.1 1686 3 483657 hornblende (j = 0.01071 ± 0.00001): a 1.8 11.847 0.123 0.04105 4 184.220 0.019 1 1 94 1965.7 11.4 b 1.9 11.642 0.027 0.00559 28.7 163.894 0.059 3.1 4.1 99.3 1828.5 3.3 •c 2 10.312 0.013 0.00193 73.5 159.127 1.198 63.2 67.3 99.9 1794.7 1.1 •d 2 10.03 0.014 0.00155 89 159.084 0.463 24.4 91.7 100 1794.4 1.3 •e 2.1 11.862 0.043 0.00182 90 158.317 0.028 1.5 93.2 100 1788.9 8.0 •f 2.6 11.62 0.018 0.00223 71.9 159.178 0.130 6.8 100 99.9 1795.1 2.3 integrated (total fusion) age: 1797 3 (•) plateau age: 95.9 1795 3 483667 hornblende (j = 0.01071 ± 0.00001): a 1.8 14 0.277 0.08890 2.2 501.205 0.011 0.5 0.5 95.1 3339.8 25.5 b 1.9 8.88 -0.010 0.01236 9.9 166.723 0.006 0.2 0.7 98.1 1848.2 27.0 c 2 8.133 0.066 0.00845 13.3 196.061 0.068 3 3.7 98.9 2041 3.3 d 2 8.823 0.082 0.01454 8.4 148.220 0.042 1.8 5.5 97.4 1715.1 6.7 e 2.1 5.161 0.073 0.01482 4.8 146.685 0.009 0.4 5.9 97.2 1703.6 17.3 •f 2.2 8.525 0.054 0.00482 24.4 157.760 0.118 5.2 11 99.3 1784.9 2.3 •g 2.2 9.17 0.048 0.00241 52.4 157.782 0.188 8.2 19.2 99.8 1785.1 1.7 •h 2.3 8.683 0.050 0.00316 37.8 157.288 0.206 9 28.2 99.6 1781.6 1.6 •i 2.3 8.979 0.056 0.00234 53 157.382 0.328 14.3 42.5 99.8 1782.2 1.4 •j 2.4 8.623 0.055 0.00105 113.4 157.274 0.432 18.8 61.3 100 1781.5 1.9 •k 2.6 8.346 0.056 0.00150 77 157.395 0.561 24.4 85.8 99.9 1782.3 1.1 •l 2.8 9.598 0.067 0.00374 35.4 156.861 0.081 3.5 89.3 99.5 1778.5 3.0 •m 4 10.4 0.062 0.00273 52.5 157.231 0.246 10.7 100 99.8 1781.1 1.9 integrated (total fusion) age: 1801 2 (•) plateau age: 94.1 1782 2 483671 muscovite (j = 0.01071 ± 0.00001): a 1.4 0.2482 0.005 0.00153 2.2 145.263 0.449 27.4 27.4 99.7 1692.9 1.2 b 1.5 0.7642 0.007 0.00189 5.6 145.703 0.168 10.3 37.6 99.6 1696.2 1.9 •c 1.5 0.3775 0.007 0.00197 2.6 143.327 0.168 10.2 47.8 99.6 1678.2 1.6 •d 1.6 0.4299 0.017 0.00377 1.6 142.856 0.089 5.4 53.3 99.2 1674.6 2.8 •e 1.6 1.5215 0.010 0.00180 11.7 142.234 0.057 3.5 56.7 99.7 1669.8 3.7 •f 1.7 0.9017 0.017 0.00144 8.7 142.377 0.093 5.7 62.4 99.7 1670.9 3.0 •g 1.7 0.2113 0.005 0.00004 79.6 143.047 0.230 14 76.4 100 1676 1.7 •h 1.8 1.123 0.015 0.00144 10.7 142.970 0.065 4 80.4 99.7 1675.4 23.0 •i 2 0.4492 0.005 0.00121 5.1 143.638 0.152 9.3 89.7 99.8 1680.5 4.2 j 2.3 0.2898 0.004 0.00172 2.3 144.357 0.169 10.3 100 99.7 1686 3.2 integrated (total fusion) age: 1684 3 (•) plateau age: 52.1 1676 3 table 1. 40ar-39ar analytical data for step heating experiments on amphiboles and muscovites from the rinkian fold belt step pwr/t°c ca/k cl/k 36ar/39ar %36ar(ca) 40*ar/39ar mol 39ar % step %39ar %40*ar age (ma) ± age cumulated (2σ) j: irradiation parameter. 40*ar: radiogenic 40ar. steps marked with dot (•) are included in the plateau age for each sample. 119 whereas the tectonic model of grocott & pulvertaft (1990) operated with four contractional and three extensional events in an epicontinental marginal basin, four main phases of deformation that developed during progressive crustal shortening are now recognised. it has been debated in recent years whether the previous distinction between the rinkian and nagssugtoqidian belts in west greenland is meaningful in tectonic terms (e.g. garde & steenfelt 1999a; van gool et al. 2002), and it has now been proposed that the two belts represent the northern and southern parts of a common, more than 1100 km wide collisional orogen, separated by a suture located in the disko bugt region (fig. 1; connelly et al. 2005). the continuous crustal shortening in the rinkian fold belt throughout its tectonic evolution is in agreement with a setting within the northern of two colliding plates at some distance from the suture, and thus in accordance with the proposed tectonic linkage to the nagssugtoqidian orogen. however, the 40ar/39ar data presented in the following section may be interpreted to indicate that the tectonometamorphic events in the rinkian and nagssugtoqidian belts were not contemporaneous. descriptions of samples and results of 40ar/39ar age determinations four hornblende samples and three muscovite samples were collected at the head of ukkusissat fjord close to the prøven igneous complex, between svartenhuk and uummannaq, and close to the north coast of nuussuaq (fig. 1). sample numbers refer to the data base of the geological survey of denmark and greenland. ukkusissat fjord near the prøven igneous complex a sample with hornblende was collected south of the prøven igneous complex, within the high-grade contact metamorphic aureole where extensive partial melting has been observed, particularly within the karrat group (grocott & pulvertaft 1990). the sample (483657, fig. 1) comes from a homogeneous, medium-grained, amphibolitic relict dyke within the regional flat-lying tonalitic orthogneiss basement. the amphibolite dyke is approximately one metre thick, a few metres long, and has been isoclinally folded. the sample is mostly composed of light to dark green hornblende between 0.5 and 1 mm in diameter, together with some plagioclase and minor phases step pwr/t°c ca/k cl/k 36ar/39ar %36ar(ca) 40*ar/39ar mol 39ar % step %39ar %40*ar age (ma) ± age cumulated (2σ) 483696 hornblende (j = 0.01071 ± 0.00001): a 1.9 4.9568 0.167 0.02251 3 557.080 0.091 1.9 1.9 98.9 3501.9 4.4 b 2 5.6395 0.198 0.00248 31.3 176.447 0.447 9.5 11.5 99.7 1914.4 1.3 c 2.1 5.7159 0.202 0.00177 44.4 164.542 0.370 7.9 19.4 99.8 1833 1.7 d 2.2 5.9588 0.197 0.00468 17.5 161.233 0.042 0.9 20.3 99.3 1809.7 6.6 e 2.2 5.7368 0.198 0.00147 53.9 161.903 0.098 2.1 22.3 99.9 1814.5 2.5 f 2.3 5.3872 0.190 0.00107 69.3 162.278 1.212 26.4 48.8 99.9 1817.1 1.0 •g 2.3 5.3337 0.184 0.00097 75.5 157.503 1.620 34.6 83.4 100 1783.1 1.3 •h 2.4 5.3052 0.209 0.00097 75.2 157.732 0.242 5.2 88.5 100 1784.7 1.6 •i 2.6 6.2547 0.255 0.00436 19.8 161.354 0.017 0.4 88.9 99.4 1810.6 13.4 •j 2.9 7.1602 0.246 0.00815 12.1 156.396 0.009 0.2 89.1 98.7 1775.1 13.0 •k 4 5.4765 0.189 0.00181 41.7 158.216 0.511 10.9 100 99.8 1788.2 1.6 integrated (total fusion) age: 1864 2 (•) plateau age: 51.2 1785 3 483708 hornblende (j = 0.01071 ± 0.00001): a 1.9 7.672 0.051 0.00834 12.7 181.053 0.134 8.2 8.2 98.8 1945 2.2 •b 2 9.5527 0.062 0.00197 66.7 157.766 0.855 52.6 60.8 99.9 1785 1.0 •c 2.1 9.6506 0.062 0.00317 42 157.637 0.376 23.2 84 99.7 1784.1 1.5 •d 2.2 8.7213 0.062 0.00050 241.2 157.711 0.054 3.3 87.3 100.1 1784.6 3.8 •e 2.3 10.597 0.109 0.00836 17.5 157.274 0.024 1.5 88.7 98.7 1781.5 8.4 •f 2.4 13.824 0.144 0.00502 38 156.401 0.012 0.7 89.5 99.4 1775.2 12.1 •g 2.7 10.368 0.067 0.00298 47.9 157.265 0.171 10.5 100 99.7 1781.4 3.3 integrated (total fusion) age: 1798 3 (•) plateau age: 91.8 1784 3 table 1 (continued) 120 such as biotite, titanite and zoizite. the biotite is intergrown with hornblende, and the plagioclase is partly altered to sericite. the obtained plateau age is 1795 ± 3 ma (fig 2; table 1). north-eastern uummannaq four samples were collected in the kangilleq–kangerlussuaq area, 75–100 km north of uummannaq (fig. 1). this area was less intensely affected by rinkian metamorphism than other areas investigated in this study, with chlorite schist locally preserved on the north coast of qeqertarsuaq. samples 483653 and 483654, both from the nukavsak formation, were collected at two localities close to each other near the southern end of qeqertarsuaq (fig. 1). sample 483653 (fig. 1) is a greywacke consisting of biotite, sillimanite, quartz, muscovite and small amounts of tourmaline and zircon. it is a fine-grained rock, where biotite and muscovite together define the main tectonic foliation. fibrolitic sillimanite occurs in broom-shaped clusters close to muscovite. it was difficult to obtain a good separate from this sample because the muscovite is very fine grained, intergrown with biotite, and sometimes has altered rims. this sample yielded a u-shaped spectrum, with a minimum which yields an age of 1681 ± 6 ma and represents 24% of the total 39ar-release (fig. 2; table 1). sample 483654 (fig. 1) is a fine-grained andalusite schist with centimetre-sized andalusite poikiloblasts in a matrix of biotite and quartz, minor tourmaline and muscovite. partial recrystallisation of andalusite to fibrolite was observed. the biotite shows two different orientations implying growth both during d2 and d3 deformation. the muscovite crystals are very small and often occur close to biotite, but sometimes also in small separate clusters. this sample gave a u-shaped spectrum, with a minimum representing 26% of the total 39ar-release and yielding an age of 1686 ± 3 ma (fig. 2; table 1). sample 483667 (fig. 1) was collected from a decimetre-thick, boudinaged, homogeneous amphibolite band in tonalitic reworked orthogneiss on qinngusaaq (fig. 1). folds, lineations, δand σ-shaped porphyroclasts and foliations representing both d2–d3 and d4 occur at the sampling locality. the light to dark brownish-green hornblende forms well-crystallised medium-grained aggregates with interstitial plagioclase partly altered to sericite. small grains of pale green pyroxene, probably diopside, occur together with the hornblende. a hornblende plateau age of 1782 ± 2 ma was obtained (fig. 2; table 1). sample 483671 (fig. 1) consists of fine-grained metasandstone to metasiltstone from the nukavsak formation with quartz, biotite, muscovite and sillimanite as main minerals. biotite, muscovite and sillimanite define the main tectonic foliation, where sillimanite often occurs in clusters containing small muscovite and biotite grains. at this locality, quartz pods display distinct asymmetries in two different directions. the asymmetric pods on rock faces with sw–ne orientations suggest top-to-ne tectonic transport (during d2), whereas rock faces with se– nw orientations suggest transport to the nw (during d3) and contain biotite lineations with that trend. muscovite from this rock, presumably grown during d3, gave a plateau age of 1676 ± 3 ma (fig. 2; table 1). south-east of uummannaq two samples with hornblende were collected from the archaean basement south-east of the uummannaq area, close to the north coast of nuussuaq (fig. 1). sample 483696 (fig. 1) comes from a hornblenditic layer in a leucogabbro that occurs as enclaves in quartzo-feldspathic orthogneiss. the sample consists almost exclusively of light to dark green, mediumto coarse-grained hornblende. the hornblende plateau age is 1785 ± 3 ma (fig. 2; table 1). sample 483708 (fig. 2) comes from an amphibolite dyke that cuts the fabric of the surrounding augen gneiss and is probably palaeoproterozoic in age. both the amphibolite and the host gneiss are intensely deformed. this sample has biotite and hornblende growing together, feldspars partly altered to sericite, and minor amounts of quartz. the hornblende plateau age is 1784 ± 3 ma (fig. 2; table 1). discussion and conclusions the results from this study provide the first published constraints on cooling ages of hornblende and muscovite in the central rinkian belt. hornblende 40ar/39ar plateau age spectra from samples 483657, 483667 and 483708 yield ages between 1795 and 1782 ma. these ages all form well-defined plateaus, and the plateaus represent more than 90% of total 39ar release. sample 483696 yielded a plateau age of 1785 ma for 51% of the total 39ar release, and is consistent with the other hornblende ages. muscovite samples 483653 and 483654 both yield u-shaped age spectra with minima representing less than 50% of the total 39ar release, at 1681 and 1686 ma respectively. sample 483671 provides a plateau age of 1676 ma defined by 121 52% of the total 39ar release. the muscovite plateau age spectrum for sample 483671 is consistent with the minima provided by samples 483653 and 483654. these taken together suggest a relatively consistent muscovite cooling age below 350°c of c. 1680 ma in the central rinkian belt. the obtained hornblende and muscovite ages at 1795– 1782 and 1686–1676 ma, respectively, are remarkably uniform, although they cover a distance of c. 200 km in chlorite to sillimanite grade amphibolite facies terrain across the entire central part of the rinkian fold belt. this 40ar/39ar age study shows that the temperatures reached during the palaeoproterozoic tectonothermal reworking were everywhere sufficiently high to reset the 40ar/39ar hornblende and muscovite systems in the rocks examined. the ages date the cooling below the closure temperature of ar diffusion in hornblende and muscovite after the palaeoproterozoic metamorphic event, and the data suggest a slow cooling rate of c. 1.5°c/ma between c. 1780 and 1680 ma, using closure temperatures of 500°c for hornblende and 350°c for muscovite (mcdougall & harrison 1999). due to recent recalculation of the primary and secondary standards used in 40ar/39ar geochronological experiments (renne et al. 1998), the previously published 40ar/ 39ar ages from the nagssugtoqidian belt (rasmussen & holm 1999; willigers et al. 2001, 2002), which use the older standard age, have to be multiplied by 1.009 in order to compare directly with the new 40ar/39ar ages presented here from the rinkian belt. in the northern part of the nagssugtoqidian orogen, willigers et al. (2001, 2002) obtained 40ar/39ar hornblende ages of 1756–1733 ma (recalculated from 1740–1717 ma) and muscovite ages of c. 1715 ma (recalculated from 1700 ma). in the central part of the orogen still farther south, their hornblende ages range between c. 1750–1700 ma and muscovite ages between c. 1765–1715 ma (recalculated from 1750–1700 ma). in the disko bugt area (fig. 1), where connelly et al. (2005) proposed a suture between the two belts, a set of 40ar/39ar and k-ar hornblende age data reported by rasmussen & holm (1999) scatter between archaean ages and a k-ar age of c. 1765 ma, revealing that temperatures during the palaeoproterozoic thermal event were not sufficiently high in all parts of this area to reset the k-ar isotope system (rasmussen & holm 1999). the uniformity of the new 40ar/39ar ages from the central and northern rinkian fold belt suggests that the ages are largely unrelated to the metamorphic grade and to the structural history of the geographical locations of the samples, with the possible exception of sample 483657 from ukkusissat fjord (see below). accordingly, the 40ar/39ar data are interpreted as regional, late orogenic cooling ages which are not directly related to the tectono-metamorphic history of the individual samples. this conclusion is supported by (in part unpublished) u-pb zircon ages of synto late-kinematic palaeoproterozoic pegmatites from the same region, which are older than 1800 ma (thrane et al. 2003; k. thrane, personal communication 2004). willigers et al. (2002) reached the same conclusion from their 40ar/39ar studies of the central and northern nagssugtoqidian orogen reported above, pointing out that their study area represents a section of middle to lower crust that was only slowly exhumed by erosion. in preserved upper crustal levels of younger orogens it is commonly possible to date specific tectonic events using the 40ar/39ar method, because the dated units were either transported rapidly to these crustal levels and are not yet eroded away, or the minerals grew at temperatures near or below their closing temperature and thus constrain the age of the prograde tectonothermal event itself. the 1795 ± 3 ma age of the hornblende from ukkusissat fjord (sample 483657, about 12 ma older than the other hornblende ages) may point to early uplift of this particular area, which is a domain of early ne-directed d2 thrusting that was not affected by the subsequent nw-directed tectonic transport during d3. the cooling rate of c. 1.5°c/ma documented by this study (using hornblende and muscovite closure temperatures of 500°c and 350°c) is only slightly slower than the 2–3°c/ma reported by willigers et al. (2001, 2002) from the central nagssugtoqidian orogen, but considerably slower than rates between 5° and 7°c/ma reported by the latter authors from the northern nagssugtoqidian orogen. willigers et al. (2001, 2002) used less accepted closure temperatures of 580°c and 410°c for hornblende and muscovite, respectively, implying a difference of 170°c between hornblende and muscovite closure temperatures. the latter temperature gap is larger than the 150°c used in this study, but this makes little difference to the calculation of cooling rates. the uniform 40ar/39ar hornblende ages resulting from the present investigation are significantly older than those in both the northern and central parts of the nagssugtoqidian orogen (willigers et al. 2001, 2002). as regards muscovite, the rinkian muscovite ages are younger than muscovite ages in the northern nagssugtoqidian belt, but older than those in the central nagssugtoqidian orogen (willigers et al. 2001, 2002). the fact that rinkian hornblende ages are older than those in the nagssugtoqidian orogen shows that cooling below 500°c took place earlier in the rinkian fold belt than in both the central and northern parts of the nagssugtoqidian orogen. it is therefore 122 plausible that uplift began significantly earlier in the rinkian belt but was slower than in the nagssugtoqidian orogen, which may in turn suggest that the main phases of compression and peak metamorphism in the two belts were not synchronous. these interpretations are consistent with the observation by taylor & kalsbeek (1990) that pb-pb whole-rock isochron ages of marbles in the two belts (interpreted as representing recrystallisation of the marbles during peak metamorphism) differ significantly from each other. marbles collected on appat island in the central rinkian belt (fig. 1) yielded a pb-pb isochron of 1881 ± 20 ma, whereas an age of 1845 ± 23 ma was obtained from marbles in the central part of the nagssugtoqidian orogen. our interpretations are also consistent with the fact that the 40ar/ 39ar data reported here show no signs of having been affected by a contact metamorphic aureole around the prøven igneous complex. the intrusion age of the latter at 1869 ± 9 ma is coeval with the youngest members of the arfersiorfik complex and sisimiut charnockite in the central nagssugtoqidian orogen (connelly et al. 2000; van gool et al. 2002). the prøven igneous complex has intruded rocks belonging to the karrat group that were already intensely deformed and metamorphosed prior to the intrusion, but before the last major deformation and peak metamorphism (thrane et al. 2005); the prøven igneous complex represents a crustal melt that was apparently not related to subduction processes. in contrast, the arfersiorfik complex and sisimiut charnockite in the south represent i-type magmas that were related to precollision subduction. notwithstanding the overall structural and geochronological evidence for a direct linkage between the rinkian fold belt and the nagssugtoqidian orogen, the age relationships outlined above may imply that collision-related deformation, metamorphism and magmatic activity took place in the northern rinkian belt while subduction was still going on south of the recently proposed suture in the disko bugt region. it may be speculated that such diachronism is also reflected in the dissimilar 40ar/39ar cooling ages from the rinkian and nagssugtoqidian parts of the entire palaeoproterozoic orogenic complex in west greenland. alternatively, the different rinkian and nagssugtoqidian cooling ages might relate to different depths of burial. however, this is not supported by the uniform hornblende 40ar/39ar cooling ages found within the rinkian belt itself, regardless of geographical distance and metamorphic facies; further discussion of large-scale platetectonic implications is beyond the scope of the present paper. analytical procedure four hornblende and three muscovite separates from the central rinkian belt have been dated with the 40ar/39armethod. the hornblende separates were obtained from amphibolite and diorite, and muscovite from metasedimentary rocks, by crushing, sieving and handpicking. the hornblende and muscovite samples selected for 40ar/39ar geochronology were irradiated together with the dra-2 sanidine standard (25.26 ma; wijbrans et al. 1995, recalculated following renne et al. 1998), for 35 hours at the nrg-petten hfr rodeo facility in petten, the netherlands. j-values (the irradiation parameter) were calculated with a precision of 0.5%. the 40ar/39ar geochronology laboratory at the university of lund employs a micromass 5400 mass spectrometer with a faraday cup and an electron multiplier. a metal extraction line, which contains two saes c50-st101 zr-al getters and a cold finger cooled to c. –155°c by a polycold p100 cryogenic refrigeration unit, is also present. one or two grains of hornblende or muscovite were loaded into a copper planchette that consists of several 3 mm holes. samples were step-heated using a defocused 50w co2 laser. sample clean-up time was 5 minutes, using the two hot zr-al saes getters and the cold finger. the laser was rastered over the samples to provide even heating of all grains. the entire analytical process is automated and runs on a macintosh computer with software developed at the berkeley geochronology center by al deino and modified for the laboratory at the university of lund. time zero regressions were fitted to data collected from 10 scans over the mass range of 40 to 36. peak heights and backgrounds were corrected for mass discrimination, isotopic decay and interfering nucleogenic ca-, k-, and cl-derived isotopes. isotopic production values for the cadmium lined position in the petten reactor are 36ar/ 37ar(ca) = 0.000270, 39ar/37ar(ca) = 0.000699, and 40ar/ 39ar(k) = 0.00183. 40ar blanks were calculated before every new sample and after every three sample steps. 40ar blanks were between 5.0 and 3 × 10–16. blank values for masses 39 to 36 were all less than 7 × 10–18. blank values were subtracted for all incremental steps from the sample signal. the laboratory was able to produce very good incremental gas splits, using a combination of increasing time at the same laser output, followed by increasing laser output. age plateaus were determined using the criteria of dalrymple & lanphere (1971), which specify the presence of at least three contiguous incremental heating steps with statistically indistinguishable ages and constituting greater than 50% of the total 39ar released during the experiment. inverse isochrons yield ages statistically indis123 tinguishable from those given by the plateaus and are not presented here. 40ar/39ar plateau age spectra are presented in fig. 2 and the analytical data in table 1. acknowledgements the authors thank j.n. connelly, j. grocott, m. hand, k.j.w. mccaffrey and k. thrane for discussions leading to the preparation of this manuscript, which also draws on their collective field observations in 2002–2003. we are grateful to j. grocott and å. johansson for critical reviews. references connelly, j.n., van gool, j.a.m. & mengel, f.c. 2000: temporal evolution of a deeply eroded orogen: the nagssugtoqidian orogen, west greenland. canadian journal of earth sciences 37, 1121–1142. connelly, j.n., thrane k., krawiec, a. & garde a.a. 2005: linking the palaeoproterozoic nagssugtoqidian and rinkian orogens through the disko bugt region of west greenland. journal of the geological society (london) 162, 1–17. dalrymple, g.b & lanphere, m.a. 1971: 40ar/39ar technique of k-ar dating: a comparison with the conventional technique. earth and planetary science letters 12, 300–308. escher, a. & pulvertaft, t.c.r. 1976: rinkian mobile belt of west greenland. in: escher, a. & watt, w.s. (eds): geology of greenland, 105–119. copenhagen: geological survey of greenland. garde, a.a. & pulvertaft, t.c.r. 1976: age relations of the precambrian mârmorilik marble formation, central west greenland. rapport grønlands geologiske undersøgelse 80, 49–53. garde, a.a. & steenfelt, a. 1999a: precambrian geology of nuussuaq and the area north-east of disko bugt, west greenland. geology of greenland survey bulletin 181, 6–40. garde, a.a. & steenfelt, a. 1999b: proterozoic tectonic overprinting of archaean gneisses in nuussuaq, west greenland. geology of greenland survey bulletin 181, 141–154. garde, a.a., grocott, j., thrane, k. & connelly, j.n. 2003: reappraisal of the rinkian fold belt in central west greenland: tectonic evolution during crustal shortening and linkage with the nagssugtoqidian orogen. geophysical research abstracts 5 (egs-agu-eug joint assembly), 09411. garde, a.a., connelly, j.n., grocott, j., hand, m., mccaffrey, k.j.w. & thrane, k. 2004: crustal shortening and granite emplacement in the rinkian fold belt, west greenland, and implications for palaeoproterozoic laurentian evolution. danmarks og grønlands geologiske undersøgelse rapport 2004/17, 16–18. grocott, j. & pulvertaft, t.c.r. 1990: the early proterozoic rinkian belt of central west greenland. in: lewry, j.f. & stauffer, m.r. (eds): the early proterozoic trans-hudson orogen of north america. geological association of canada special paper 37, 443–463. henderson, g. & pulvertaft, t.c.r. 1987: geological map of greenland, 1:100 000, mârmorilik 71 v.2 syd, nûgâtsiaq 71 v.2 nord, pangnertôq 72 v.2 syd. descriptive text, 72 pp., 8 plates. copenhagen: geological survey of greenland. kalsbeek, f. 1981: the northward extent of the archaean basement of greenland – a review of rb-sr whole-rock ages. precambrian research 14, 203–219. kalsbeek, f., pulvertaft, t.c.r. & nutman, a.p. 1998: geochemistry, age and origin of metagreywackes from the palaeoproterozoic karrat group, rinkian belt, west greenland. precambrian research 91, 383–399. mcdougall, i. & harrison, t.m. 1999: geochronology and thermochronology by the 40ar/39ar method. oxford: oxford university press. pulvertaft, t.c.r. 1986: the development of thin thrust sheets and basement-cover sandwiches in the southern part of the rinkian belt, umanak district, west greenland. rapport grønlands geologiske undersøgelse 128, 75–87. ramberg, h. 1949: on the petrogenesis of the gneiss complexes between sukkertoppen and christianshaab, west greenland. meddelelser dansk geologisk forening 11, 312–327. rasmussen, h. & holm, p.m. 1999: proterozoic thermal activity in the archaean basement of disko bugt region and eastern nuussuaq, west greenland: evidence from 40ar–39ar mineral age investigations. geology of greenland survey bulletin 181, 55–64. renne, p.r., swisher, c.c., deino, a.l., karner, d.b., owens, t.l. & depaolo, d.j. 1998: intercalibration of standards, absolute ages and uncertainties in 40ar/39ar dating. chemical geology 145, 117–152. taylor, p.n. & kalsbeek, f. 1990: dating the metamorphism of precambrian marbles: examples from proterozoic mobile belts in greenland. chemical geology (isotope geoscience section) 86, 21–28. thrane, k., connelly, j.n., garde, a.a., grocott, j. & krawiec, a.w. 2003: linking the palaeoproterozoic rinkian and nagssugtoqidian belts of central west greenland: implications of new u-pb and pbpb zircon ages. geophysical research abstracts 5 (egs-agu-eug joint assembly), 09275 only. thrane, k., baker, j., connelly, j.[n.] & nutman, a.p. 2005: age, petrogenesis and metamorphism of the syn-collisional prøven igneous complex, west greenland. contributions to mineralogy and petrology 149, 541–555. van gool, j.a.m., connelly, j.n., marker, m. & mengel, f.c. 2002: the nagssugtoqidian orogen of west greenland: tectonic evolution and regional correlation from a west greenland perspective. canadian journal of earth sciences 39, 665–686. wijbrans, j.r., pringle, m.s., koppers, a.a.p & scheveers, r. 1995: argon geochronology of small samples using the vulkaan argon laserprobe. proceedings of the koninklijke nederlandse akademie van wetenschappen 98, 185–218. willigers, b.j.a., krogstad, e.j. & wijbrans, j.r. 2001: comparison of thermochronometers in a slowly cooled granulite terrain: nagssugtoqidian orogen, west greenland. journal of petrology 42, 1729– 1749. willigers, b.j.a., van gool, j.a.m., wijbrans, j.r., krogstad, e.j. & mezger, k. 2002: posttectonic cooling of the nagssugtoqidian orogen and a comparison of contrasting cooling histories in precambrian and phanerozoic orogens. journal of geology 110(5), 503–517. _________________________________________________________________________________________________________________________________________________________________________________ manuscript received 21 june 2004; revision accepted 15 march 2005 124 research article gregersen et al. 2025: geus bulletin 60. 8385. https://doi.org/10.34194/299dt488 1 of 29 seismic investigations of eight geological structures for potential storage of co2 in denmark: an introduction ulrik gregersen1* , michael b.w. fyhn1 , marie keiding1 , tanni abramovitz1 , morten bjerager1 , henrik vosgerau1 , florian w.h. smit2 , thomas funck1 , anders mathiesen1 , finn mørk2 , niels h. schovsbo2 , henrik i. petersen2 , lars henrik nielsen1 , karen dybkjær2 , bodil w. lauridsen2 , emma sheldon2 , maiken l. olsen3 , gunver k. pedersen1 , carsten m. nielsen2 , erik s. rasmussen1 , alireza malehmir4, myrto papadopoulou4 , samuel zappalá4 , magdalena markovic4 , michael westgate4 , jolanta putnaite4 , emmanouil konstantinidis4 , kristina kucinskaite4 , axel ehrhardt5 , egon nørmark6 1department of geophysics and sedimentary basins, geological survey of denmark and greenland (geus), copenhagen, denmark, 2department of geo-energy and storage, geological survey of denmark and greenland (geus), copenhagen, denmark, 3department of geosciences and natural resource management, university of copenhagen, copenhagen, denmark, 4department of earth sciences; geophysics, geocentrum, uppsala university, uppsala, sweden, 5subdepartment marine resource exploration, federal institute for geosciences and natural resources (bgr), hannover, germany, 6department of geoscience, aarhus university, aarhus, denmark abstract in june 2021, a novel danish national carbon capture and storage strategy was ratified by the danish parliament, and this was followed by the initiation of the project ‘ccs2022–2024’, led by the geological survey of denmark and greenland. in collaboration with other institutions, we acquired and interpreted new 2d seismic data between 2022 to 2024 to investigate and mature eight sites for potential subsurface storage of co2 in danish onshore and offshore areas. this bulletin contains a series of papers that present important results of the work. in this introduction paper, we provide an overview of seismic acquisitions and the interpretation of seismic data together with existing deep wells. the study sites selected are large subsurface structures located in onshore jylland, sjælland and lolland and offshore denmark in the eastern north sea. the onshore targets are the gassum, havnsø, rødby, stenlille and thorning structures, while the offshore sites comprise the inez, jammerbugt and lisa structures. the project work comprises a series of reports regarding extensive seismic acquisition, processing and interpretation of the new and pre-existing seismic data as well as other publications emanating from the project. this bulletin and the technical reports present an improved understanding of the formation, composition and geometry of the investigated structures. the studies include the mapping of the reservoir and seal formations, identification of principal faults, interpretation of the stratigraphic and structural development, reservoir and seal characterisation and estimates of the static storage capacity. hence, this research provides a significant step forward concerning characterisation of the geology and maturation of the potential storage sites. in addition, it has inspired new ideas, including an updated regional stratigraphic interpretation of the triassic succession of the danish basin and correlation with adjacent basins. *correspondence: ug@geus.dk received: 27 nov 2024 revised: 04 mar 2025 accepted: 06 03 2025 published: 21 may 2025 keywords: carbon capture and storage (ccs), co2 storage, denmark, gassum formation, seismic acquisition abbreviations: bmsl: below mean sea level bgr: german federal institute for geosciences and natural resources ccs: carbon capture and storage db: danish basin ecu: early cimmerian unconformity fm: formation geus: geological survey of denmark and greenland gp: group grv: gross rock volume mcu: mid-cimmerian unconformity mems: micro electronic mechanical systems n/g: average ratio of net sand to gross reservoir volume ngb: north german basin obs: ocean bottom seismometers rfh: ringkøbing–fyn high rts: realtimeseismic sc: storage capacity stz: sorgenfrei–tornquist zone twt: two-way travel time geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: jon r ineson (geological survey of denmark and greenland, denmark) reviewed by: daniel sopher (geological survey of sweden), alvar braathen (university of oslo, norway) funding: see page 26 competing interests: none declared additional files: none provided https://doi.org/10.34194/299dt488 https://orcid.org/0000-0001-5946-2582 https://orcid.org/0000-0003-4039-5847 https://orcid.org/0000-0002-2933-0199 https://orcid.org/0009-0001-9241-5937 https://orcid.org/0000-0003-3180-8857 https://orcid.org/0000-0002-7582-5360 https://orcid.org/0000-0001-8438-386x https://orcid.org/0000-0003-3646-490x https://orcid.org/0000-0003-4345-7513 https://orcid.org/0000-0002-3134-6835 https://orcid.org/0000-0003-4723-0586 https://orcid.org/0000-0001-6606-7062 https://orcid.org/0000-0002-2078-7360 https://orcid.org/0000-0002-8420-3379 https://orcid.org/0000-0002-4060-746x https://orcid.org/0000-0003-4353-8241 https://orcid.org/0000-0002-2042-3655 https://orcid.org/0000-0002-0792-2257 https://orcid.org/0000-0002-1525-1385 https://orcid.org/0000-0001-8603-8429 https://orcid.org/0000-0002-5112-2732 https://orcid.org/0000-0003-3402-6850 https://orcid.org/0000-0003-2902-7349 https://orcid.org/0000-0002-4199-7439 https://orcid.org/0000-0002-2215-3097 https://orcid.org/0009-0002-1475-2230 https://orcid.org/0009-0007-8687-0929 https://orcid.org/0000-0001-9842-0940 https://orcid.org/0000-0001-5472-2736 mailto:ug@geus.dk gregersen et al. 2025: geus bulletin 60. 8385. https://doi.org/10.34194/299dt488 2 of 29 geusbulletin.org 1. introduction the need for reduction of greenhouse gas emissions to the atmosphere to counter further climate deterioration is becoming increasingly urgent. carbon capture and storage (ccs) is an important strategy for considerably lowering atmospheric co2 emissions (ipcc 2005), and is currently implemented in a growing number of countries since the first projects were initiated more than 25 years ago, for example at the sleipner field offshore norway (the sacs project; chadwick et al. 2004; gregersen & johannessen 2007). in the danish sector of the north sea, the first pilot geological storage of co2 started in march 2023 (the greensand project; szabados & poulsen 2023). the danish subsurface has been considered highly suitable for geological co2 storage for many years, and screening studies document a large geological storage potential that is widely distributed onshore and offshore (e.g. larsen et al. 2003; anthonsen et al. 2014; hjelm et al. 2022). detailed and site-specific studies based on thorough interpretation of new subsurface data from eight selected structures (fig. 1) conducted by the ccs2022– 2024 project (gregersen et al. 2023a, 2023b; abramovitz et al. 2024; bjerager et al. 2024; fyhn et al. 2024; keiding et al. 2024) suggest a reduction in expected static storage capacity compared to previous estimates (hjelm et  al. 2022). the significant danish co2 storage potential is based on the favourable geology that includes good reservoir properties of regionally distributed reservoirs, thick seals, large structures and relatively quiescent present tectonic activity. the storage potential is contained within sandstone reservoirs (saline aquifers), and the danish onshore and nearshore areas contain several structures with significant co2 storage potential. eight structures were investigated in the ccs2022–2024 project led by the geological survey of denmark and greenland (geus). five sites are onshore structures (the gassum, havnsø, rødby, stenlille and thorning structures), and three are offshore structures (the inez, jammerbugt and lisa structures; fig. 1). new interpretations of seismic and well data were made for the eight structures. for six of the structures, new seismic surveys were acquired, whereas for the lisa and inez structures sufficient pre-existing data were available. the focus has been on integrating the new seismic data with the fig. 1 map of the danish structures with potential for geological storage of co2 from the geus-led ccs2022–2024 project. the dark green structures (stenlille, havnsø, rødby, gassum, thorning, jammerbugt, lisa and inez structures) are mapped in the project. outlines of the lisa and inez structures are modified from hjelm et al. (2022). the dark green shading shows the extent of the deepest mapped closure of the top gassum fm surface, except in the rødby structure, where the deepest closure of the top bunter sandstone fm is delineated. the light green shading shows outlines of other structures, which may have potential for geological storage of co2. 100 km 4ºe 58ºn 58ºn 57ºn 56ºn 55ºn 6ºe 8ºe 10ºe 12ºe inez lisa jammerbugt gassum thorning havnsø stenlille rødby structures selected for maturation 2022–2024 legend other mapped structures n fyn sjælland falster møn lolland jylland https://doi.org/10.34194/299dt488 gregersen et al. 2025: geus bulletin 60. 8385. https://doi.org/10.34194/299dt488 3 of 29 geusbulletin.org previously acquired seismic and well data. the six seismic surveys were acquired in 2022 and 2023. the first seismic survey was acquired in 2022 at the ne flank of the stenlille structure as a small test survey and was upscaled for the other onshore surveys (e.g. papadopoulou et al. 2023, 2024). uppsala university was contracted by geus for the onshore seismic acquisition. the german federal institute for geosciences and natural resources (bgr) and aarhus university performed the marine seismic acquisition of the jammerbugt structure for geus. for the havnsø structure, aarhus university also acquired data in the marine strait between sjælland and nekselø, recording signals from the vibro-truck sources onshore near the coast. consultancy firm cowi assisted with the logistics, applications and communication, while students from copenhagen, uppsala, aarhus and unilasalle universities supported the seismic field work. the scope of this special issue of geus bulletin is to provide an overview of the many results of the ccs2022– 2024 project, including the context, purpose and methods that have improved the geological understanding and maturation of the eight selected structures for potential co2 storage. the bulletin contains ten individual research papers. this first introductory paper includes the context, aim and summary of methods and key results of the project, and provides a common reference for the remaining papers of the bulletin. six papers focus on the investigated structures. the inez, gassum, rødby and thorning structures are described in individual papers. however, due to their proximity and similar geological development, the stenlille and havnsø structures are grouped and described together in a single paper, as are the jammerbugt and lisa structures. three papers are topic-specific in which common elements of the structures such as the reservoir properties, their potential static storage capacity, brine composition and seal capacity are described and discussed. 2. geological setting the danish basin trends wnw−ese between the ringkøbing–fyn high to the south and the sorgenfrei– tornquist zone and the skagerrak–kattegat platform to the north (figs 2, 3). the north german basin is situated south of the ringkøbing–fyn high. the danish basin is the eastern part of the larger norwegian–danish basin extending from the norwegian north sea in the northwest to sjælland in the east. the danish basin is an intracratonic basin that developed since the late palaeozoic when it was initiated by late carboniferous – early permian crustal extension reflected in normal faulting associated with widespread magmatism (ziegler 1990; michelsen & nielsen 1991, 1993; vejbæk 1997; abramovitz et al. 1998, 2000). pre-dating the danish basin, the oldest documented sedimentary successions resting on crystalline basement are lower palaeozoic rocks known from a few deep wells and from outcrops and wells on bornholm fig. 2 top pre-zechstein map of the main structural elements onshore and offshore denmark, including highs, basins and main faults. the elements include the norwegian–danish basin, the eastern part of which is named the danish basin, the sorgenfrei–tornquist zone, the skagerrak–kattegat platform, the ringkøbing–fyn high and the northern part of the north german basin. modified from vejbæk (1997). 100 km 4ºe 58ºn 57ºn 56ºn 55ºn 6ºe 8ºe 10ºe 12ºe 14ºe 16ºe >9 km 8–9 km 7–8 km 6–7 km 3–4 km 5–6 km 4–5 km 2–3 km 1–2 km 0–1 km structural high fault zone boundary salt pillow salt diapir n o r w e g i a n – d a n i s h b a s i n n o r w e g i a n – d a n i s h b a s i n r i n g k ø b i n g – f y n h i g h east north sea high farsund basin egersundbasin rønne graben fjerritslev trough s o r g e n f r e i – t o r n q u i s t z o n e s k a g e r r a k k a t t e g a t p l a t f o r m himmerland graben sweden nor way germany top pre-zechstein ø resund basin east sjæ lland h igh n o r t h g e r m a n b a s i n brande trough h orn g raben central g raben https://doi.org/10.34194/299dt488 gregersen et al. 2025: geus bulletin 60. 8385. https://doi.org/10.34194/299dt488 4 of 29 geusbulletin.org and in kattegat, and adjacent regions in southern sweden and norway (nielsen & japsen 1991; erlström et al. 1997; erlström & sivhed 2012; schovsbo et al. 2016; nielsen & klitten 2023). the oldest sedimentary rocks comprise cambrian sandstones and mudstones representing continental and near-shore deposition in a shallow epicontinental sea. later cambrian and ordovician deposition of organic-rich mud and carbonates resulted from continued transgression. subsequently, a thick interval of organic-rich silurian shale was deposited within a deepening foreland basin (nielsen & schovsbo 2011, 2015; schovsbo et al. 2016). the danish basin is located north of the e–w-striking caledonian deformation front straddling southern denmark, along which the east avalonia microcontinent collided with the baltica plate (including scandinavia) during the ordovician to early devonian caledonian orogeny (ziegler 1990; babel working group 1993; abramovitz et al. 1998). during the caledonian orogeny, northwards movements of east avalonia resulted in thrusting and deformation of the lower palaeozoic succession in se denmark and the sw baltic sea (lassen et al. 2001). the norwegian–danish and north german basins started to form in response to lithospheric stretching and rifting in the carboniferous−permian. the ringkøbing–fyn high, located between the two basins (fig. 2), evolved during the same period, forming broad horsts with lesser extension compared with the adjacent basins (vejbæk 1997). extensional faulting was associated with volcanism and led to the formation of large, rotated fault blocks, extensive erosion, and widespread, mostly coarse siliciclastic deposition (carboniferous and overlying rotliegend group; gp; michelsen & nielsen 1991, 1993; vejbæk 1997; stemmerik et al. 2000; nielsen 2003). in some places, the rotliegend gp forms discrete syn-rift wedges on rotated fault blocks (michelsen & nielsen 1991; vejbæk 1997). the region farther to the east (bornholm and southern sweden) forms part of the major sorgenfrei–tornquist zone (fig. 2), where complex strike-slip tectonism and pull-apart basins evolved (erlström et al. 1997; vejbæk 1997). the top pre-zechstein surface (top of the rotliegend gp where the group is preserved, see fig. 3) is in part a significant unconformity and represents one of the deepest regionally mappable levels from seismic data onshore and offshore denmark. this surface outlines the main structural elements described above (figs 2, 3; vejbæk 1997). in the late permian (zechstein), late-rift thermal subsidence dominated, and major basins with restricted seaway connections developed. the northern and southern permian basins (see outlines in peryt et al. 2010) were partially separated by the ringkøbing– fyn high and are mainly characterised by deposition of evaporites and carbonates included in the zechstein gp (fig. 3; stemmerik & frykman 1989; stemmerik et al. 2000; peryt et al. 2010). following the permian, regional subsidence continued, and a thick triassic siliciclastic-dominated succession accumulated, comprising sandstones and mudstones with subordinate carbonates and evaporite intervals. the triassic lithostratigraphic subdivision adopted in this study (bunter shale, bunter sandstone, ørslev, falster, tønder, oddesund and vinding formations) follows bertelsen (1978, 1980). the triassic climate was warm and with some exceptions arid. deposition of the bunter shale and bunter sandstone formations (fms) occurred in fluvio-limnic and continental dominated environments with desert sand plains and sabkhas during deposition of the bunter sandstone fm (bertelsen 1980; clemmensen 1986; bachmann et al. 2010). at the same time, the skagerrak fm developed farther to the north reflecting the fluvially-alluvially dominated environment that bordered the scandinavian craton (olsen 1988). during the early triassic, especially in southern denmark, deposition occurred in large lakes, sabkhas, playas and maybe even short-lived shallow seas (ørslev fm, equivalent to röt fm in the north german basin, see fig. 3; bertelsen 1980). meanwhile, deposition farther to the north remained continental in character. the connection to the tethys sea and the north european epicontinental sea in the south increased during the anisian (early middle triassic) and mudstones and carbonates were deposited. limestone beds are frequent in the falster fm, which is equivalent to the muschelkalk fm in the north german basin (fig. 3; bertelsen 1980; lindström et al. 2017). marine-influenced interludes can be traced far into the norwegian–danish basin that was otherwise dominated by continental deposition (michelsen & clausen 2002). during the ladinian (late middle triassic), the epicontinental sea retreated to the south and coastal plains and playas developed (bertelsen 1980). later, deposits became more sand-prone, influenced by fluvial processes and interbedded with finer-grained sediments. the nature of the sedimentary regime and the presence of plant remains in the upper tønder fm, and time-equivalent deposits in the danish basin and scania are indicative of slightly more humid conditions at this time (bertelsen 1980; lindström et al. 2017). lateral thickness variations and faulting within the tønder fm in southern denmark suggest active tectonism and extension along the northern margin of the north german basin. salt mobilisation occurred simultaneously with the extension. following the carnian (early late triassic), tectonism and uplift resulted in the establishment of the early cimmerian unconformity https://doi.org/10.34194/299dt488 gregersen et al. 2025: geus bulletin 60. 8385. https://doi.org/10.34194/299dt488 5 of 29 geusbulletin.org fig. 3 schematic stratigraphic diagram oriented from north to south including wells in the sorgenfrei–tornquist zone (stz), the eastern part of the danish basin (db), ringkøbing–fyn high (rfh) and the north german basin (ngb), compiled in this study. the well sections are from north to south located geographically in the kattegat sea (terne-1), and on the islands of sjælland (stenlille-19, slagelse-1), fyn (ullerslev-1), falster (ørslev-1) and lolland (søllested-1, rødby-2) shown by red vertical lines (arrows indicate where well termination, total depth, is located below the figure). well locations are shown in fig. 4. the lithostratigraphy is based on available well sections (nielsen & japsen 1991 and references therein) and new research from this project using wells and seismic data. similar groups and formations occur in jylland, and the figure is used as a reference for the current danish lithostratigraphy (incl. bertelsen 1978, 1980; michelsen et al. 2003; nielsen 2003). the upper permian-triassic lithostratigraphy of northern germany is also shown. modified from abramovitz et al. (2024). röt fm bunter sandstone fm bunter shale fm zechstein kupferschiefer muschelkalk fm keuper fm 145 150 155 160 165 170 175 180 185 190 195 200 140 135 80 85 90 95 100 105 110 115 120 125 130 75 70 260 265 240 245 250 255 235 230 220 225 210 215 205 u pp er ju ra ss ic m id dl e ju ra ss ic lo w er ju ra ss ic tr ia ss ic pe rm ia n lo w er c re ta ce ou s u pp er c re ta ce ou s oxfordian kimmeridgian volgian callovian bathonian bajocian aalenian toarcian pliensbachian sinemurian hettangian rhaetian norian carnian ladinian anisian olenekian induan changhsingian capitanan wordian wuchiapingian maastrichtian campanian santonian coniacian turonian cenomanian albian aptian barremian hauterivian valanginian ryazanian skagerrak fm bunter sandstone fm bunter shale fm chalk gp rødby fm vedsted fm børglum fm flyvbjerg fm fjerritslev fm mcu ecu gassum fm vinding fm oddesund fm falster fm ørslev fm zechstein gp rotliegend gp haldager sand fm frederikshavn fm tønder fm hiatus arid lacustrine, restricted marine mudstones, evaporites and thin sandstones fluvial and aeolian sandstones and mudstones marine marlstones and mudstones marine limestones and mudstones evaporites and mudstones shallow marine and �uvial sandstones marine mudstones marine chalk 65 paleocene danian ngb rfh db stz ns rø db y2 sø lle st ed -1 ø rs le v1 sl ag el se -1 st en lil le -1 9 u lle rs le v1 t e rn e1 u. permian – triassic lithostratigraphy of northern germany mcu: mid-cimmerian unconformity ecu: early cimmerian unconformity https://doi.org/10.34194/299dt488 gregersen et al. 2025: geus bulletin 60. 8385. https://doi.org/10.34194/299dt488 6 of 29 geusbulletin.org (ecu) described in the north german basin and over the ringkøbing–fyn high (clausen & pedersen 1999; ahlrichs et al. 2020), but presumably also existing over part of the danish basin (fig. 3). a significant hiatus characterised by the absence of large parts of the carnian is recorded in southern danish wells in the north german basin (e.g. lolland, falster and southern jylland; fig. 3). during the late triassic, deposition of mudstones and evaporites, mainly in the middle of the oddesund fm (fig. 3) indicate a return to mainly arid conditions, separated by episodes of more humid conditions. in some places, sand was also deposited (bertelsen 1980). in the danish basin, mobilisation of zechstein salt into salt pillows was initiated regionally during the deposition of the oddesund fm due to differential loading, deep-seated tectonism and faulting (boldreel 1985; geil 1991). the oddesund fm therefore shows large variations in thickness throughout much of the danish basin governed by salt migration, rifting and differential subsidence across the basin. thickness variations are most distinct towards nw jylland, in the fjerritslev trough and within the himmerland graben (fig. 2). some of the salt pillows evolved into salt diapirs and other types of salt structures (e.g. salt walls) in the north german basin, the western danish basin (nw jylland) and farther west in the norwegian–danish basin (north sea; boldreel 1985; sørensen 1998). during the uppermost late triassic (late norian to rhaetian), more humid conditions were established, and the danish basin became marine-influenced (bertelsen 1980). mud-dominated sediments with calcareous and sandy interbeds (vinding fm) were deposited in shallow, brackish-marine environments in the deeper parts of the basin, while sand-dominated deposition was initiated at the margins (fig. 3; bertelsen 1980). during the latest triassic (rhaetian) and into the earliest jurassic (hettangian – early sinemurian), coastal to continental areas were repeatedly overstepped by the sea, and fluvial, coastal and shallow marine sand interbedded with offshore mud were deposited, which now constitute the widely distributed gassum fm (bertelsen 1978; nielsen 2003). the gassum fm is the key reservoir formation for potential co2 storage within the structures located in the danish basin, which are described in this bulletin. continued rise in relative sea level, on a regional scale, during the early jurassic resulted in widespread deposition of thick clay-dominated successions with more silty and sandy interludes (fjerritslev fm). the gassum and fjerritslev fms have been subdivided sequence-stratigraphically into sequences and systems tracts, which can be correlated throughout the basin (nielsen 2003). middle−late jurassic regional uplift related to the mid-cimmerian tectonic phase led to major erosion in large parts of the danish basin (mid-cimmerian unconformity, see fig. 3), in places enhanced by vertical salt movement. the hiatus expands towards the ringkøbing– fyn high and the north german basin (figs 2, 3; nielsen 2003). the middle jurassic to earliest cretaceous tectonism and uplift probably contributed to sand-rich deposition (haldager sand and frederikshavn fms), mainly from the north across the skagerrak–kattegat platform and sorgenfrei–tornquist zone, and southwards into the danish basin (figs 2, 3). pauses in tectonism and renewed subsidence during the late jurassic to early cretaceous caused increased clayand mud-rich deposition (flyvbjerg and børglum fms; nielsen 2003). mud-dominated deposition (vedsted fm) continued during the early cretaceous (valanginian to aptian) and became more calcareous with marl and chalk units during the albian (rødby fm). upper cretaceous chalk was deposited throughout the danish basin followed by danian limestone, together constituting the chalk gp (fig. 3). s–nand sw–ne-oriented alpine orogeny compression led to episodes of late cretaceous to paleocene regional inversion, which affected large parts of northern europe. inversion tectonism also impacted the danish basin and the sorgenfrei–tornquist zone (ziegler 1990; vejbæk 1997). in the early part of the cenozoic, deposition in the danish basin was influenced by post-rift thermal subsidence combined with the opening of the north atlantic ocean. the hemipelagic sediments consist of fully marine clay and marl and, locally, diatomite and ash layers (ziegler 1990; heilmann-clausen 1995; schiøler et al. 2007 and references therein). subsequent uplift of the danish basin occurred during the late oligocene and early miocene, and at the same time parts of fennoscandia were inverted (ziegler 1990; japsen & bidstrup 1999; japsen et al. 2007; rasmussen et al. 2008, 2010). this resulted in the development of large sand-rich fluvio-deltaic systems in present-day jylland. glacioeustatic sea-level changes also influenced the sedimentation pattern. also, during the miocene many salt structures were active in the central graben and most likely also in the norwegian–danish basin. after the early miocene, a new tectonic regime formed (rasmussen 2009). during the middle miocene, accelerated subsidence of the north sea basin, including the danish area, occurred. the delta plains were flooded, and the fluvio-deltaic successions were overlain by fully marine mud (rasmussen et al. 2010). mud deposition continued during the late miocene, although uplift was initiated in the late tortonian. deposits from the pliocene are not present onshore denmark. https://doi.org/10.34194/299dt488 gregersen et al. 2025: geus bulletin 60. 8385. https://doi.org/10.34194/299dt488 7 of 29 geusbulletin.org uplift and tilting towards the sw occurred during the quaternary, resulting in erosion of older deposits (rasmussen et al. 2005). hence, in the eastern part of the danish area, quaternary deposits overlie upper cretaceous chalk and danian limestone, while in the southwestern part of jylland, the youngest pre-quaternary deposits are of tortonian (late miocene) age (sorgenfrei & bertelsen 1954; rasmussen et al. 2010). processes during the quaternary, predominantly associated with glaciations and de-glaciations, deformed and altered large areas of denmark (houmark-nielsen 1987, 2004). 3. database and methods 3.1. database an important part of the ccs2022–2024 project was the acquisition of new seismic data to improve the understanding of the potential for co2 storage within the eight selected structures (fig. 1). the pre-existing database of the study includes wells (fig. 4) and seismic data (fig. 5) of the structures. vintage seismic profiles provide the basis for mapping structures selected for maturation (figs 5, 6). a review of vintage seismic data and of the structures was used for planning of new seismic profiles of the structures for updated mapping (maturation) in this project. the acquisitions of new seismic surveys (figs 5, 6) were carried out during 2022–2023 and were focused over the six structures: gassum, havnsø, jammerbugt, rødby, stenlille and thorning (figs 1, 5). a few seismic lines were also acquired over the lisa structure to tie to the jammerbugt structure, but apart from these lines, sufficient legacy seismic lines cover the lisa and inez structures to map these areas. in addition, deep wells in nearby structures (fig. 4) were also used to tie well data (lithology, formations, ages etc.) and to consider reservoir and seal properties. in total, 2093 km of 2d seismic profiles were acquired in the project. a total of 643 km profiles were acquired fig. 4 locations of danish wells (red circles) and outlines of regional structural elements, including structural highs, basins, fault zones and faults. modified from nielsen (2003). 10°e 12°e 12°e8°e 14°e 56°n 55°n skåne h or n fyn sjælland jylland fyn lolland falster møn g ra be n bornholm sorgenfrei–tornquist zone danish basin ringkøbing – skagerrak– kattegat platform sweden north german basin germany high fjerritslev trough børglum fault fjerritslev fault haldager fault east sjæ lland h igh øresund basin structural elements in denmark high fault well border 50 km 8°enor way 58°n 57°n 56°n 55°n horsens-1 voldum-1 rønde-1 terne-1 gassum-1 hobro-1 vemb-1 mejrup-1 rødding-1 farsø-1 hyllebjerg-1 års-1 vedsted-1 vendsyssel-1 frederikshavn-1, -2, -3 flyvbjerg1 sæby-1 thisted-2,-3 nøvling-1 grindsted-1 mors-1 oddesund-1 skive-1 skive-2 børglum-1 haldager-1 skagen-2 slagelse-1 lavø-1 fjerritslev-1,-2 inez-1 k-1 jelling-1 brøns-1 margretheholm-1,-2 karlebo-1 kværs-1 hønning-1 tønder wells rødekro-1 åbenrå varnæs-1 glamsbjerg-1 kegnæs-1 sønderborg-1,-2 rødby-1,-2 ringe-1 ullerslev-1 stenlille-19 stenlille wells stenlille-1 søllested-1 ørslev-1 erslev-1,-2 linde-1 arnum-1 løve-1 harte-2 uglev-1 feldsted-1 j-1 c-1 felicia-1 thisted-1, -4 erik-1 borg-1 løgumkloster-1, -2 kvols-1 vinding-1 hans-1 anholt-4 pernille-1 stina-1 https://doi.org/10.34194/299dt488 gregersen et al. 2025: geus bulletin 60. 8385. https://doi.org/10.34194/299dt488 8 of 29 geusbulletin.org over the onshore structures (including a few kilometres in the marine strait at havnsø) while 1450 km were acquired offshore, centred over the jammerbugt structure (table 1). examples of vintage seismic data and new seismic data of this project from onshore and offshore areas are shown in fig. 6. 3.2. onshore seismic acquisition the existing seismic database over the selected onshore structures consisted mostly of sparse and poor-quality seismic data acquired during hydrocarbon exploration between the 1960s and the 1980s (e.g. fig. 6a, c). these older data typically have relatively low vertical resolution, a poor signal-to-noise ratio and discontinuous seismic reflections affected by noise. this means that stratigraphic and structural details are often uncertain and not clearly observable in these data compared to more recent data (fig. 6). furthermore, the coverage of these older data is often low (large and irregular spacing between profiles); only a few legacy seismic lines are available over the thorning, gassum, havnsø structures, the central to eastern rødby structure, and the north-eastern flank of the stenlille structure. a 3d seismic survey from 1997 and more recent 2d lines were available over the central part of the stenlille structure. as a result, new 2d seismic profiles were acquired in some areas over these structures. geus contracted uppsala university (sweden) to acquire and process the five onshore seismic surveys at the stenlille, havnsø, rødby, gassum and thorning structures. the seismic source was provided by geopartner geofizyka (poland). cowi coordinated permissions and logistics. the surveys had field assistance by students from copenhagen, aarhus, and unilasalle universities. the seismic data were collected along roads using vibroseis trucks and a dual recording system (fig. 7). two 12t vibroseis trucks generated the seismic source at shot points with 10 m intervals, with a peak force of 95 kn and synchronised sweeps with a linear frequency increasing from 10 hz to 140 hz in 18 s. at every shotpoint location, this sweep was repeated three times to improve the signal-to-noise ratio in the processing. the record length was set to 25 s after each shot. on average, 2.5–3.0 km of seismic data were acquired per day. the dual recording system consisted of a landstreamer towed behind the rear truck and wireless geophones placed along the profile (fig. 7). it is designed and fig. 5 locations of the new seismic lines (red) acquired in 2022 and 2023, and outlines (green) of the structures that were mapped as part of this project (ccs2022–2024). older seismic lines are shown in blue. 100 km 4ºe 58ºn 57ºn 56ºn 55ºn 6ºe 8ºe 10ºe 12ºe rødby new seismic lines acquired in the ccs2022–2024 project legend structures selected for maturation 2022–2024 existing seismic lines n inez lisa jammerbugt gassum thorning havnsø stenlille fyn sjælland falster møn lolland jylland https://doi.org/10.34194/299dt488 gregersen et al. 2025: geus bulletin 60. 8385. https://doi.org/10.34194/299dt488 9 of 29 geusbulletin.org tested to produce two complementary high-resolution data sets for an improved imaging of both the shallow and the deeper subsurface (malehmir et al. 2022; zappalá et al. 2022). the seismove® landstreamer, developed by uppsala university, has micro electronic mechanical systems (mems) sensors mounted 2 m apart and with a 1 ms sampling interval. the landstreamer is comprised of 40 m segments that are attached end-to-end, providing a flexible length of 40 to 240 m. adjustment of the streamer length depended on the logistical complexity of the given day’s traverse due to road bends, buildings and traffic. the wireless geophones were placed along roads at 10 m intervals (nodal system), and they have a natural frequency of 10 hz with a 2 ms sampling interval. fig. 6 examples of vintage and newly acquired 2d seismic data shown with vertical scale in two-way travel (twt) time in milliseconds (ms). example from onshore seismic data across the havnsø structure with (a) vintage seismic data (ref: ssl6267-r12) and (b) new seismic data acquired and processed by uppsala university (ref: geus22-hvn-p7). example from offshore seismic data across the jammerbugt structure with (c) vintage seismic data (ref: wgc64a-39970) and (d) newly acquired seismic data by the federal institute for geosciences and natural resources (bgr) and aarhus university, and that was subsequently reprocessed by realtimeseismic (rts; ref: geus23-jb-15 from the reprocessed survey geus2023-jammerbugt-re2023). –1000 –2000 –3000 –1000 –2000 –3000 –1000 –2000 –3000 –1000 –2000 –3000 sw sw nene tw t ( m s) tw t ( m s) tw t ( m s) ns sw ne 10 kmjylland jammerbugt d c jammerbugt structure tw t ( m s) havnsø structure sjælland havnsø sejerø bugt 10 km (a) (b) (c) (d) a b jammerbugt structure havnsø structure n 0 2 km n 0 2 km 0 2 km 0 2 km https://doi.org/10.34194/299dt488 gregersen et al. 2025: geus bulletin 60. 8385. https://doi.org/10.34194/299dt488 10 of 29 geusbulletin.org geophone positions were derived using differential gps measurements with an accuracy of 10–30 cm. both the operational zone and the active spread for each recording system changed each day during the acquisition, with mostly a 200 m spread length for the landstreamer, and typically 5–9 km spread length for the nodal system. each seismic profile was independently processed; the overall processing flow was kept consistent, however, with only minor variations in the input parameters per profile. the processing steps are thoroughly described in the acquisition and processing reports (malehmir & papadopoulou 2022, 2023; malehmir & westgate 2023; malehmir & markovic 2024; putnaite & malehmir 2024). the landstreamer and nodal data were processed independently, yielding final stacked and migrated sections. furthermore, the data sets were merged in the pre-stack domain, rebalanced and jointly stacked and migrated to produce a third section per profile. merging the wireless geophone data and the landstreamer data unites each data set configuration’s benefits and optimises the merged data set’s signal-to-noise ratio (e.g. papadopoulou et al. 2022; malehmir et al. 2025). in addition, the data were reprocessed by realtimeseismic (2023a, b; 2024a, b, c, d) to further enhance the resolution of the data and ensure optimal ties between the seismic lines. particular effort was put into the reprocessing to understand and suppress crooked-line artefacts caused by the irregular source locations along roads with bends (abramovitz et al. 2024). table 1 seismic data (2d) acquired 2022 and 2023 during the ccs2022–2024 project and pre-exisiting data. geological structure line acquired pre-existing wells and seismic data stenlille structure 13 km (5 profiles) 20 stenlille wells, one 3d seismic survey (1997) and c. 20 seismic profiles (of 2d seismic surveys from 1962–1967, 1972–73, 1981, 1987, 1994) havnsø structure 131 km (9 profiles) no wells, c. 20 seismic profiles (of 2d seismic surveys from 1962–1967, 1972–73, 1974) gassum structure 259 km (14 profiles) gassum-1 well, 6 seismic profiles (of 2d seismic surveys from 1962–1967, 1973–74, 1981–1983) rødby structure 106 km (12 profiles) rødby-1 & -2 wells, c. 12 seismic profiles (of 2d seismic surveys from 1962–1967, 1979–1981) & a few more recent lines from a shallow seismic survey thorning structure 134 km (8 profiles) no wells, 2 seismic profiles (of a 2d seismic survey from 1973–74) acquired onshore 643 km acquired offshore for the jammerbugt structure 1450 km (39 profiles) no wells, 3 seismic profiles (of 2d surveys from 1964, 1982), partly over the structure total acquired 2093 km fig. 7 photographs from the seismic acquisition with the two vibroseis trucks, landstreamer and geophones. (a) operational setup of field equipment. (b) mems sensors mounted at 2 m intervals on a landstreamer towed behind the rear vibroseis truck. (c) two vibroseis trucks are operated with synchronised vibrations. (d) wireless geophones are deployed every 10 m along the profile. reproduced from malehmir & westgate (2023). fo ur to s ix s tr ea m er s eg m en ts (2 0 un it s pe r s eg m en t) aa bb cc dd 2m 10m https://doi.org/10.34194/299dt488 gregersen et al. 2025: geus bulletin 60. 8385. https://doi.org/10.34194/299dt488 11 of 29 geusbulletin.org 3.3. offshore seismic acquisition 3.3.1. havnsø–nekselø seismic acquisition to map the northern extent of the havnsø structure, the onshore seismic survey included a profile (geus22hvn-p1) that extends offshore (fig. 8; funck & nørmark 2023; malehmir & papadopoulou 2023). this line took advantage of the nearby island of nekselø that is separated from sjælland by a 2 km wide shallow marine strait. due to the shallow water depth and the protection status as a natura2000 area, no seismic sources could be employed offshore. instead, seismic receivers were deployed in the marine strait to record the onshore vibroseis sweeps both on sjælland and nekselø. while shallow reflectivity cannot be traced continuously across the marine strait because of the marine source restrictions, deeper structures that are relevant for the assessment of the havnsø structure can be mapped from sjælland to nekselø. for the offshore recording, 18 ocean bottom seismometers (obs) of the type sercel microbs were deployed at the seafloor between havnsø and nekselø in august 2022 (fig. 8). these instruments are equipped with three-component geophones and a hydrophone; the latter provided better data quality in this experiment. the obs recorded the seismic signals that were initiated on the road towards havnsø, and subsequently on nekselø. some obs stopped recording prior to the completion of the line. however, the data quality of the recorded signals is good. after correlation with the source sweep, reflections can be seen from depths below 2 s two-way-travel time (twt). a second type of receiver was used close to havnsø, where a 600 m long marine streamer with 96 channels was deployed from the shore and seaward (fig. 8). additional weights kept the streamer at the seafloor. motion transferred from the recovery buoy resulted in poorer data at the tail end of the streamer. a complete account of the marine acquisition component and the initial processing of the data is provided by funck & nørmark (2023), and malehmir & papadopoulou (2023) describe the merging with the land data. 3.3.2. jammerbugt seismic acquisition prior to this study, the seismic data coverage of the jammerbugt structure was very limited, where only three pre-existing seismic lines were available. hence, the acquisition of additional data to allow for a proper assessment of the structure’s suitability for underground co2 storage was required. therefore, 1450 km of 2d seismic data were acquired in april 2023 covering most of the jammerbugt structure in a denser line spacing (c. 2–3 km) and systematic grid than existed previously (figs 5, 6, 9). the survey was carried out using the faroese research vessel jákup sverri (fig. 10) in a collaboration between geus, aarhus university fig. 8 overview of the havnsø seismic acquisition from havnsø to nekselø. (a) location map showing the onshore seismic sources (vibroseis) and the marine receivers (obs: ocean bottom seismometers, and streamer cable) used along the northern portion of seismic line geus22-hvn-p1. (b) schematic cross-section of the marine receivers. (c) the streamer winch on the beach at havnsø with 600 m of streamer cable deployed at the seafloor towards nekselø. (photograph by per trinhammer 2022) (d) four of the retrieved obs on board a local fishing vessel that was used for the marine operation. (photograph by egon nørmark 2022). n (b) water winch obs obs obs beach havnsønekselø streamer (600 m) 3 8 (a) ocean bottom seismometer (obs) end points streamer cable streamer cable survey lines geus22-hvn-p1 exisiting 2d seismic lines (pre 2022) geus22 – hvn–p1 sjælland havnsø nekselø august 24, 2022 seism ic sources august 23, 2022 sejerø bugt 12 10 6 9 5 7 4 14 20 15 16 1718 19 11 13 obs (d) nekselø winch streamer (c) seism ic sources 1000 m https://doi.org/10.34194/299dt488 gregersen et al. 2025: geus bulletin 60. 8385. https://doi.org/10.34194/299dt488 12 of 29 geusbulletin.org and the bgr in germany. data acquisition was limited to water depths greater than 10 m, and a 10 km distance had to be kept from the natura2000 marine protected areas. the design of the survey ensured ties to the existing legacy seismic grid and to the j-1 well on the lisa structure. thirty-nine seismic lines (geus23-jb-1 to 39; fig.  9) were acquired using a 2100 m long sercel sentinel ssrd streamer (fig. 10) with 336 channels and a group spacing of 6.25 m. the seismic source was changed after acquisition of the first ten lines to provide better penetration into the hard subsurface with shallow chalk. initially, two 150 cubic inch (2.4 l) gi guns were used at a shot rate of 6 s. later, two gi guns with a volume of 355 cubic inch (5.8 l) each were employed for which a shot rate between 10 and 12 s could be maintained. the pressure was 135 bar (13.5 mpa). more details on the acquisition can be found in funck et al. (2023). fig. 9 topographic map with the location of the acquired seismic lines of the geus23-jb survey (red lines). yellow circles indicate the positions of wells, the black line marks the track of the acquisition vessel jákup sverri, blue lines indicate the 10 m depth contour. fig. 10 photographs from the jammerbugt seismic acquisition. (a) the faroese research vessel jákup sverri about to leave hirtshals harbour (northern jylland) for the seismic acquisition. (b) deployment of the streamer tail buoy. in the foreground, one of the two winches with the solid-state streamer cable (sercel sentinel ssrd; yellow cable). photographs: thomas funck. https://doi.org/10.34194/299dt488 gregersen et al. 2025: geus bulletin 60. 8385. https://doi.org/10.34194/299dt488 13 of 29 geusbulletin.org 3.4. seismic interpretation and mapping within the project the geological development and stratigraphy of each of the eight structures were investigated and evaluated, with a focus on the reservoir-seal pairs most important for co2 storage. to do this, we used wells and new and legacy seismic data. petreltm software was used for establishing the database with wells and seismic data and for wellto-seismic ties, seismic interpretation and maps. key wells in and close to structures were examined, and we mainly used the lithostratigraphy (formation tops) from nielsen & japsen (1991). if needed, the well-log lithology and ages of successions from biostratigraphy were revised, and the resulting updated well-log figures were reported (see e.g. gregersen et al. 2023a, appendix b). to use well-log data and well tops (depth domain) for the interpretation of the seismic data (time domain), seismic-well tie procedure was performed on wells that contained sonic and density logs. we used well-to-seismic ties with synthetic seismograms to relate geology to seismic response, and time-to-depth relationships in the wells were mostly from nielsen & japsen (1991; see more details e.g. in gregersen et al. 2023a). the most significant reflections were tied from wells to seismic sections and include for most structures near to base and near to top (here simply base and top) of formations or groups: top and base chalk (gp), top fjerritslev (fm), top and base gassum (fm), top bunter sandstone (fm), top zechstein (gp), and top pre-zechstein. ‘fm’ or ‘gp’ are omitted in the seismic horizon names. a strong amplitude reflection with the associated acoustic impedance change (velocity and density decrease) at the base chalk gp (e.g. fig. 6b: c. 900 ms) was used as a key for interpretation of seismic sections and is indicative for the polarity of the data. seismic reflections, seismic successions and seismic facies were identified and interpreted (e.g. gregersen et al. 2022, 2023a; smit et al. 2022). the workflow for this process was to start by interpreting the most significant seismic reflections and relate these to well-tied formation tops. the well-to-seismic ties with synthetic seismograms were used to more exactly select a trough or a peak reflection that corresponded to the formation tops. then the auto-track function was specified for each horizon to be interpreted. if the auto-tracking could not be used (e.g. where reflections were discontinuous or weak), then interpretation was performed manually. different seismic data displays were used for interpretation. the displays include various colour scales (black-white and different colour displays including black-white-red) and different seismic attributes. seismic stratigraphic relationships such as onlaps, downlaps and truncations were used to interpret seismic units and boundaries (e.g. unconformities). faults, salt structures and folds were also identified and mapped together with thickness patterns. horizon flattening of seismic profiles, faults and maps were used for the structural and tectonostratigraphic interpretation. mapping and interpretation focused particularly on the horizons and faults related to the successions of the primary reservoir (gassum fm or bunter sandstone fm) and their primary seals (fjerritslev fm or ørslev fm, respectively). in addition, the gassum fm is described in more detail in studies of sequence stratigraphy applying well-logs and seismic lines to better predict reservoir properties. reservoir properties, mapped reservoir formation thicknesses and top reservoir closure areas are included for calculations of static storage capacity of co2. methods for preparing time-to-depth conversion of maps include the compilation of a regional velocity model constructed to convert the interpreted horizons from the time domain to the depth domain. the model area was defined so that the velocity model included a significant buffer around the structure. the data used were: (1) twt seismic horizons of the main stratigraphic units, using the seismic survey lines, gridded to 250 × 250 m and adjusted to the wells in the structure; (2) well top markers; and (3) seismic migration (root mean square, rms) velocities from the 2d lines, which were converted to average velocities using the dix formula. more details are available in gregersen et al. (2023a) and abramovitz et al. (2024). 3.5. investigation of reservoir and seal the geology of the reservoir and seal successions are described in the individual, structure-specific project reports (e.g. gregersen et al. 2023a; abramovitz et al. 2024; keiding et al. 2024) and briefly in the bulletin papers of this volume using well completion reports, publications and in-house studies of well logs and geological samples from wells (mainly from cores). in addition, some studies focusing on description of lithology and biostratigraphy are described (e.g. gregersen et al. 2023a; abramovitz et al. 2024; keiding et al. 2024). the data used are from the wells in or closest to the relevant structure. the aim of these studies is to provide a more detailed understanding of the reservoir and seal characteristics for each structure. reservoir characteristics are derived mainly from acquired wireline logs, that are calibrated against conventional core analysis, descriptions of cuttings and sidewall cores. similarly, the wireline logs are used to estimate thicknesses and identify mudstone sealing sections. the specific method of the seal characterisation depends on the availability of logs in each well. key reservoir parameters are also investigated and used for storage capacity calculations. https://doi.org/10.34194/299dt488 gregersen et al. 2025: geus bulletin 60. 8385. https://doi.org/10.34194/299dt488 14 of 29 geusbulletin.org 3.6. storage capacity to compare the potential co2 storage of the structures, we use a simple equation for saline aquifers (e.g. goodman et al. 2011), where static theoretical storage capacity of reservoir units with buoyant trapping is estimated by: sc = grv * n/g * φ * ρco2r * seff (1) sc is the storage capacity in mass of co2 (mt). grv is based on seismic interpretation and depth conversion, and is the gross rock volume confined by the upper and lower boundaries of the gross reservoir interval, where thickness and outline area of the structure is defined by the top point depth and the deepest closing contour depth constrained by the spill-point depth (fig. 11). n/g is the average ratio of net sand to gross reservoir volume for the aquifer being investigated. this is based on the petrophysical and geological understanding of the thicknesses derived from the nearest wells and transformed into structure-specific geological-based average values. ɸ is the average effective reservoir porosity of the reservoir within the grv, and ρco2r is the average co2 density at reservoir pressure and temperature where the density is estimated using the ‘calculation of thermodynamic state variables of carbon dioxide’ web tool essentially based on span & wagner (1996), and wischnewski (2007). for this calculation, we assume hydrostatic pressure and geothermal gradients between 27°c/km and 30°c/km (fuchs et al. 2020). the storage efficiency factor (seff) relates to the fraction of the total available pore volume that can store co2 within the grv. this fraction depends on many subsurface aspects including the size of storage domain, heterogeneity of the formation, compartmentalisation, permeability, porosity, pressure increase, temperature, salinity and compressibility, but is also strongly influenced by different well configurations, injection schemes and displacement efficiency (e.g. wang et al. 2013). evaluation and estimation of the static co2 storage capacity in deep saline aquifers is complex, and accurate estimations of storage capacity are only reasonable at local site-specific scales. the estimated co2 storage capacity is the maximum amount that theoretically can be injected until it reaches the boundaries (i.e. deepest closing contour; fig. 11). estimation of co2 storage capacity is uncertain due to a lack of knowledge on the storage efficiency factor that is used to reduce the storage capacity to a more realistic estimation (e.g. bachu et al. 2007; hall 2008; gorecki et al. 2009; goodman et al. 2011). the stenlille structure is the best-known case in the danish onshore area, and in previous studies, a fixed storage efficiency factor of 0.4 was used (e.g. hjelm et al. 2022). this value represents a geologically excellent and well-described gassum formation sandstone reservoir with a well-defined four-way dip closure without significant cross-cutting faults offsetting the reservoir and overlying seal. for all other potential structures and reservoirs in the danish onshore area, a reduced factor of 0.1 is selected based on limited well data, geological understanding and the presence of cross-cutting faults. to address uncertainty related to the reservoir data, depth conversion, reservoir thickness estimates and co2 density ranges, simple monte carlo simulation was carried out (fig. 11). even though this methodology is simple, the purpose is to assess and illustrate the variation of the estimated co2 storage capacities as a supplement to the mean calculated values. the method is used in all structures of the ccs2022–2024 project. estimation of storage capacity assumes a static approach where the pores in the trap are expected to be 100% connected. it does not include dynamic pressure build-up, solubility of co2 in brine, co2 mineralisation reactions, presence of salt with possible movement of co2 and in-place brine (water) in the saline aquifer, neither inside nor outside the trap. detailed dynamic reservoir simulation must take these factors into account and will obviously produce different and more realistic co2 storage capacity results than those estimated with the static method used fig. 11 conceptual profile (a–a’) across a potential structure. the uncertainty in mapping the structure results in the hypothetical minimum (min.) and maximum (max.) scenarios that are very different from the most likely mapped scenario. variance in area and in gross thickness (t) will affect the gross rock volume (grv) of the structure. the uncertainty is addressed by applying uncertainty on the resulting grv and other parameters and by conducting simple monte carlo simulation to calculate 90, 50 and 10% percentiles (e.g. burruss et al. 2009; heidug et al. 2013). map of structural uncertainty on extent well c-1x well a-1x max. most likely min. a’ a n 20 km areal extent max. max. most likely min seal spill point (min) spill point (ml) spill point (max.) a a’ https://doi.org/10.34194/299dt488 gregersen et al. 2025: geus bulletin 60. 8385. https://doi.org/10.34194/299dt488 15 of 29 geusbulletin.org here – see schovsbo et al. (2025, this volume) for further discussion. a more detailed realistic dynamic reservoir simulation including well design and injection strategy is normally carried out by the awarded license holders and operators. 4. results of seismic interpretation and storage capacity estimation the seismic interpretation carried out in the ccs2022–2024 project used both the newly acquired and legacy seismic data, as well as information from deep wells. the project focused on initial maturation and de-risking by mapping and describing mainly the reservoir and seal formations, the largest faults and the geometry of structural closure, and by outlining the stratigraphic and structural development of the structures, as summarised below. 4.1. stratigraphy, geometry and geological development of individual structures 4.1.1. stenlille structure the stenlille structure is a four-way dip closure, elongated in a sw–ne direction, located in the eastern part of the danish basin, in central-west sjælland (figs 1, 12a). the main influence on the structure is a salt pillow, which predominantly impacts the dome-shaped overlying triassic and jurassic successions and slightly affects the cretaceous succession (fig. 12a). the salt pillow was mapped, described and defined by gregersen et al. (2023a), and is subsequently referred to as the stenlille salt pillow. the structure is penetrated by 20 wells and covered by both 2d and 3d seismic data, thereby offering the most comprehensive database of all the investigated structures. thus, only a small survey with five lines was acquired in 2022 on the ne flank, just outside the area covered by the 3d seismic survey. the new data increased the overall coverage of modern, high-quality data in this part of the structure (gregersen et al. 2022, 2023a; papadopoulou et al. 2022, 2023; fig. 5). through the work in this study, a greatly improved understanding of the structure has been achieved, particularly concerning the understanding of the primary reservoir-seal pair of the rhaetian gassum fm and the uppermost rhaetian to lower jurassic fjerritslev fm (gregersen et al. 2023a). a geological cross section of the structure is shown in fig. 12a. deep below the gassum fm reservoirs and mudstones are secondary reservoirs consisting of triassic sandstones of the oddesund fm (carnian to norian) and the bunter sandstone fm (olenekian to anisian). the gassum fm consists of six reservoir zones with sandstones. details on reservoir properties and geophysical modelling of these zones are given by bredesen et al. (2022, 2023) and gregersen et al. (2023a). the zones have good reservoir properties with the deepest zones (zones 5 and 6) showing the highest average effective reservoir porosities (approximately 25 and 27%, respectively), which decrease slightly in the upper zones (zones 1–4; gregersen et al. 2023a). natural gas is stored by gas storage denmark a/s within the gassum fm of the stenlille structure and extracted for consumers as needed. the storage of gas has been operated safely for more than 30 years, proving the caprock integrity of the stenlille structure. the uppermost reservoir sandstones of the gassum fm in the stenlille structure are overlain by a 240–300 m thick section of mudstones of the fjerritslev fm. the gassum fm is on average 150 m (140–160 m) thick within the structure with the top at c. 1450 m below mean sea level (bmsl). the deepest closure of the gassum fm top surface is at 1475 m bmsl, outlining an area of 5.4 km2 (gregersen et al. 2023a). the stenlille structure mainly evolved by the growth of a salt pillow, which led to the formation of an anticlinal dome structure in the overlying strata. the structure developed predominantly during the burial of the thick fjerritslev fm. the salt pillow overlies pre-zechstein successions that form the base of the present stenlille structure. the pre-zechstein successions were mainly faulted during the carboniferous – early permian, and some of the faults were probably reactivated during the triassic to jurassic. renewed faulting probably triggered salt migration over the inclined pre-zechstein succession (near top of the rotliegend gp) leading to the formation of the domal salt pillow in the stenlille structure (fig. 12a). continued growth of the salt pillow elevated the overburden, particularly during the late early jurassic to earliest cretaceous. sedimentary successions of these ages are missing in the stenlille wells (fig. 3; gregersen et al. 2023a). normal faults and, to a lesser extent, minor reverse faults, are observed in the stenlille data. the faults were manually interpreted as well as outlined by machine learning, thereby increasing the understanding of the 3d fault network (gregersen et al. 2020, 2022, 2023a; lorentzen et al. 2022). the faults in the gassum–fjerritslev levels show ne−sw trends. the top fjerritslev fm unconformity (mid-cimmerian unconformity, mcu; fig   3) is onlapped by a succession including the vedsted and rødby fms, which are overlain by the chalk gp (figs 3, 12a). shallower faults in the chalk gp have predominant strikes in three directions: nw−se, wnw−ese and ne− sw, and may be related to renewed episodes of uplift. 4.1.2. havnsø structure the havnsø structure is a four-way dip closure, elongated in a se–nw direction, located in the eastern part of the danish basin, in western sjælland (figs 1, 2, https://doi.org/10.34194/299dt488 gregersen et al. 2025: geus bulletin 60. 8385. https://doi.org/10.34194/299dt488 16 of 29 geusbulletin.org 12b); new seismic data were acquired over the structure in 2022. the deep, central part of the havnsø structure has recently been described, mapped and defined as a large salt pillow – the havnsø salt pillow – based on the new and legacy seismic data (gregersen et al. 2023b). development of the salt pillow caused the doming of the overlying triassic and jurassic successions (fig. 12b). fig. 12 simplified geological cross sections through the eight structures investigated in this study. the approximate locations of the largest faults are shown. the primary reservoir formation is the gassum fm for all structures except for the rødby structure (c), where it the primary reservoir formation is the bunter sandstone fm. (a) stenlille structure. (b) havnsø structure. (c) rødby structure. (d) gassum structure. (e) thorning structure. (f) jammerbugt structure. (g) lisa structure. (h) inez structure. note that the sections have different vertical and horizontal scales. continues on next page. stenlille-19 km2 3 1 2 0 d ep th (k m ) od vi fa, ør stenlille salt pillow sw ne bs ga fj (a) stenlille structure havnsø structure (havnsø)(kalundborg) nesw havnsø salt pillow vi od fa, ør bs ga fj (b) 2 km d ep th ( km ) 0 1 2 4 5 3 0 d ep th (k m ) ns rødby-2 rødby salt pillow fa od ør vi bs fj ga km2 (c) rødby structure 1 2 d ep th (k m ) gassum-1 sn gassum salt pillow ør tø, favi, o d fr, bø fl, hs sk ga fj 0 2 km (d) 2 1 4 5 gassum structure large fault well bunter shale fm bunter sandstone (bs) fm post chalk group chalk group rødby & vedsted (ve) fms frederikshavn (fr), børglum (bø), flyvbjerg (fl), haldager sand (hs) fms gassum (ga) fm vinding (vi), oddesund (od), skagerrak (sk) fm zechstein group pre-zechstein fjerritslev (fj) fm tønder (tø), falster (fa), ørslev (ør) fms https://doi.org/10.34194/299dt488 gregersen et al. 2025: geus bulletin 60. 8385. https://doi.org/10.34194/299dt488 17 of 29 geusbulletin.org no wells have been drilled in the structure, and the new seismic data acquired in 2022 (fig. 5) are thus tied to the deepest well in the nearby stenlille structure (the stenlille-19 well, see fig. 3). the seismic acquisition also included the crossing of a marine strait to connect with the small island nekselø (figs 4, 6; section 3.3.1). the interpretation of the havnsø structure and its reservoir and seal formations were largely based on correlation to the stenlille structure (gregersen et al. 2023b; papadopoulou et al. 2023, 2024; zappalá et al. 2024). hence, further investigations are required to address uncertainties with the seal, faults, near surface geology and velocity models (gregersen et al. 2023b; kucinskaite et al. 2023). the primary reservoir-seal pair in both the stenlille and havnsø structures comprises the gassum and fjerritslev fms (gregersen et al. 2023b). a geological cross section through the havnsø structure is shown in figure  12b. deeper-lying sandstones of the oddesund fm and the bunter sandstone fm, known from the stenlille structure, may form secondary reservoirs in the havnsø structure. the gassum fm is on average 150 m thick (with variations between 130 and 170 m) and has its top at 1550 m bmsl. the deepest closure is at 1710 m bmsl, outlining an area of 70 km2 (gregersen et al. 2023b). the reservoir properties are mainly prognosed from the nearby stenlille wells, where three scenarios are considered: (1) the reservoir properties are directly extrapolated from the nearby stenlille wells; (2) average reservoir property values from stenlille wells (stenlille-1 & -19) and selected surrounding wells are assumed; and (3) the reservoir properties are extrapolated from the nearby stenlille wells, taking into account fig. 12 (continued) simplified geological cross sections through the eight structures investigated in this study. the approximate locations of the largest faults are shown. the primary reservoir formation is the gassum fm for all structures except for the rødby structure (c), where it the primary reservoir formation is the bunter sandstone fm. (a) stenlille structure. (b) havnsø structure. (c) rødby structure. (d) gassum structure. (e) thorning structure. (f) jammerbugt structure. (g) lisa structure. (h) inez structure. note that the sections have different vertical and horizontal scales. inez-1 j-1n s 0 1 2 3 4 5 2km 6 0 1 2km 2 3 4 5 6 7 8 0 1 2 3 4 5 6 7 9 10 jammerbugt structure inez structure lisa structurethorning structure nw se sn 2km 8 9 (h) (g) (f) skod od sk ga,vi,od ga,vi fj ga,vi sk fj fj hsve, fr, bø, fl ve, fr, bø, fl ve, fr, bø, fl hs hs 5 4 3 2 1 0 2km ga vi, od fj tø fa ør sk thorning salt pillow ‘mound’ (fj) fr d ep th (k m ) d ep th (k m ) d ep th (k m ) d ep th (k m ) w e (e) https://doi.org/10.34194/299dt488 gregersen et al. 2025: geus bulletin 60. 8385. https://doi.org/10.34194/299dt488 18 of 29 geusbulletin.org the stratigraphy inferred from the seismic interpretation. the interpretation indicates that mainly sequences 4 through 6 with sand-prone lowstand system tracts reached the havnsø area, whereas upper sequences of the gassum fm may be more mudstone-dominated. this interpretation is supported by acoustic impedance modelling and differentiated seismic velocities (gregersen et al. 2023b). the interpretation of the available data suggests that good reservoir properties can be expected in all three scenarios. the havnsø structure formed in a similar way to the stenlille structure and is also underlain by a salt pillow; salt movement was initiated during the triassic. this most likely occurred after deposition of the falster fm and at the same time as the deposition of the lower part of the oddesund fm (possibly during the early–middle carnian). at this time, a significant succession of sand was deposited (intra oddesund sandstone beds observed in stenlille-19 well; gregersen et al. 2023a, 2023b). these sands were probably associated with the onset of tectonism that culminated in the development of an erosional unconformity. this unconformity can probably be correlated to an unconformity and a significant hiatus (carnian age) recognised farther to the south, close to the top of the tønder fm in the søllested-1 well. this event is probably equivalent to the early cimmerian unconformity (ecu; fig. 3), known more regionally from, for example, the north german basin (ahlrichs et al. 2020). the salt pillow mainly developed during jurassic to early cretaceous times, when burial by a thick sedimentary succession and episodic tectonic activity probably triggered salt migration. nearly all successions overlying the salt pillow form four-way dip closures. the salt pillow is underlain by a slightly inclined pre-zechstein succession (rotliegend gp and older), which is faulted. new faulting and reactivation of pre-existing faults probably occurred during the triassic and jurassic and may have triggered salt migration. there are apparently no large faults through the mesozoic in the havnsø structure, based on the present seismic database (gregersen et al. 2023b). however, the newly acquired 2d and legacy seismic surveys have a relatively wide line spacing, and a new, denser data set (preferentially a 3d survey) is needed to investigate whether critical faults are present in the structure. a denser grid would provide possibilities for more accurate evaluation of the extent, throw and pathways of faults in the seal and near-surface succession for de-risking, but would also result in improved facies analysis to predict reservoirs. minor faults in the mesozoic succession are recognised in the data and were developed during the growth of the salt pillow. the throw on these faults is typically small (mostly less than 15 ms), and the lateral extent of the faults is typically not more than a few kilometres. the faults typically trend nw−se and sw−ne, parallel to the flanks of the structure. thinning of the fjerritslev fm over the top of the structure (fig. 12b) indicate elevation of the structure and associated erosion during late early jurassic – early cretaceous. onlap of the lower cretaceous vedsted fm on the top fjerritslev horizon marks a major hiatus related to the mcu (see fig. 3), as observed in the stenlille structure. 4.1.3. rødby structure the rødby structure is a four-way dip closure, elongated in a nw–se direction, near the town of rødby in the south of the island of lolland (figs 1, 12c). the main influence on the structure is the rødby salt pillow which gives rise to the dome-shaped geometry of the overlying triassic and jurassic successions (fig. 12c; abramovitz et al. 2024). the primary reservoir-seal pair in the rødby structure is represented by the bunter sandstone fm and the mudstone seal successions of the ørslev and falster fms, which are shown in figs 3 and 12c. the reservoir and seal successions in the rødby structure are intersected by the rødby-1 and rødby-2 wells drilled in the 1950s (figs 3, 4). legacy seismic data exist in the central and western part of the structure. new data were acquired over the central part and the eastern flank of the structure (fig. 5). these new data significantly improved the data coverage and formed the basis for improved mapping and definition of the structure (abramovitz et al. 2024; malehmir et al. 2025). the bunter sandstone fm is 200–260 m (on average 230 m) thick and contains three separate sandstone-dominated intervals. the formation top is at 1100 m bmsl with the deepest closure at 1415 m bmsl, outlining an area of 117 km2 (abramovitz et al. 2024). the rødby-1 and rødby-2 wells show fair to good reservoir properties within the bunter sandstone fm, with average porosities of 17–32% and permeabilities of 293–2029 md (abramovitz et al. 2024). the domed-shaped anticlinal structure evolved sequentially over at least three main episodes of salt migration (abramovitz et al. 2024), namely: (1) initial growth of the rødby salt pillow, which began during deposition of the falster and tønder fms as indicated by truncations along the southern flank. a thick sediment cover resulted in increased pressure and temperatures, which initiated movement of the salt, whereas faulting together with the inclined pre-zechstein succession influenced the resultant movement (top rotliegend gp; fig. 12c). (2) growth of the salt pillow during the middle jurassic to early cretaceous, triggered by tectonism and re-activated faulting during the mid-cimmerian tectonic https://doi.org/10.34194/299dt488 gregersen et al. 2025: geus bulletin 60. 8385. https://doi.org/10.34194/299dt488 19 of 29 geusbulletin.org phase resulting in thickness variations of the fjerritslev fm. salt pillow growth elevated the structure and was associated with erosion of the fjerritslev fm and perhaps later deposits. the top fjerritslev surface is equivalent to the mcu, separating the lower jurassic fjerritslev fm from the lower cretaceous vedsted fm (fig. 3). this tectonism may also have formed the small faults at the top of the structure (fig. 12c) slightly offsetting parts of the fjerritslev fm and older successions. (3) final growth of the salt pillow and faulting after deposition of the paleocene, probably up until recent times (hiatus until quaternary), slightly doming the landscape above the structure. quaternary (glacial) valleys formed, in particular, north of the rødby structure along zones of weakness that have nearly the same orientation as the deeper-seated faults (lolland–falster fault zone, see abramovitz et al. 2024). the lolland–falster fault zone mainly developed during the late triassic and was reactivated during middle jurassic and cenozoic times. the structural evolution leading to the outline of the present-day rødby structure may have been triggered by reactivation of the deep-seated palaeozoic faults below the top pre-zechstein at the base of the present-day rødby structure. 4.1.4. gassum structure the gassum structure is a four-way dip closure, elongated in an e–w direction, located centrally in the danish basin (figs 1, 2), in eastern jylland (fig. 1). the main influence on the structure is the gassum salt pillow, mapped and defined using the new seismic data by keiding et al. (2024), which gives rise to the domal geometry in the overlying triassic and jurassic successions (fig. 12d). the reservoirs and seal successions of the gassum structure are intersected by the gassum-1 well, located centrally in the structure (fig. 12d). other nearby wells (e.g. hobro-1 and voldum-1; fig. 4) are included in the interpretation to add further information on the geology of the structure. only a few legacy seismic lines of poor quality covered the structure before the new survey was acquired in 2023 (fig. 5). the newly acquired data provide much-improved data coverage for updated mapping (westgate et al. 2023, 2024, 2025; keiding et al. 2024) but also underline the need for further investigation of the seal, overburden and faults for site characterisation and risk assessment (konstantinidis et al. 2023; keiding et al. 2024). the primary reservoir-seal pair in the gassum structure consists of the gassum and fjerritslev fms (fig. 12d; keiding et al. 2024). a deeper-situated, secondary reservoir-seal pair consists of the sandstone-rich skagerrak fm overlain by the lower to lowermost upper triassic mudstone seal succession of the ørslev–falster–tønder fms. the ages of these formations are shown in fig. 3. in addition, a secondary, shallower reservoir-seal pair consists of the uppermost upper jurassic to lower cretaceous frederikshavn and lower cretaceous vedsted fms. although not present in the gassum-1 well, the middle jurassic haldager sand and upper jurassic flyvbjerg fms (fig. 3) possibly form additional reservoirs in the structure, as these seem to be present and thicken downdip at the northern flank of the structure, where they are overlain by a potential seal of the børglum fm mudstones and overlying successions (fig. 12d). the secondary reservoirs may provide a significant upside to the estimated storage capacity of the gassum fm in the gassum structure. the gassum fm is c. 130–180 m thick in the structure and is overlain by a more than 300 m thick mudstone succession referred to the fjerritslev fm. the shallowest point of the gassum fm within the structure is 1375 m bmsl and the lowermost closing contour is at 2300 m bmsl, outlining an area of 280 km2 (keiding et al. 2024). the gassum-1 well shows good reservoir properties for the gassum fm sandstones with an average effective reservoir porosity of 28.5% and an average permeability of 1500 md (keiding et al. 2024). the uniform thickness of the skagerrak fm infers no or only minor syn-depositional faulting and salt tectonism. this also applies to the overlying triassic succession (keiding et al. 2024). most of the triassic succession, including the gassum fm, is intersected by e–w-trending faults, but faulting does not significantly affect the internal thicknesses. consequently, faulting is interpreted to have occurred mainly after deposition of the fjerritslev fm. minor movements during the early jurassic, however, are indicated by thickening of the fjerritslev fm towards the north. salt movements and uplift leading to development of the gassum structure may have started during the triassic, but renewed movement and uplift occurred after deposition of the fjerritslev fm. during these periods, parts of the middle to late jurassic haldager sand and flyvbjerg fms may have been eroded or faulted out since they are not recorded in the gassum-1 well (fig. 12d; keiding et al. 2024). it is possible that these formations occur as part of a wedge, which thins towards the northern top of the structure (fig. 12d). this implies that the gassum structure probably has secondary reservoir potential; the haldager sand and flyvbjerg fms are known from other wells (e.g. the hobro-1 well to the west; fig. 4). the lower cretaceous vedsted fm forms a seal for the frederikshavn fm and thins towards the top of the structure (fig. 12d). the rødby fm and most of the vedsted fm are missing in the gassum-1 well, thereby revealing a major hiatus in the upper part of the https://doi.org/10.34194/299dt488 gregersen et al. 2025: geus bulletin 60. 8385. https://doi.org/10.34194/299dt488 20 of 29 geusbulletin.org structure and indicating uplift and erosion, likely related to salt movements (keiding et al. 2024). the prevailing e–w-striking faults offset the jurassic to uppermost cretaceous successions but can be recognised at all stratigraphic levels from the top of the ørslev fm to the chalk gp (fig. 12d). large-scale faults that pass through the seal and continue up to the shallow succession are observed. these faults can be a critical risk for co2 storage and must be addressed in further studies, which should include dense seismic acquisition and risk assessment. additional investigation of the leakage risks associated with the old well should also be conducted. the faulting is interpreted to have been associated with salt migration, and consequently, the youngest salt movements must be late cretaceous or cenozoic in age. the decreasing chalk thickness over the crest of the gassum structure may have been caused by uplift and erosion, which has been ascribed to salt movement or structural inversion (konstantinidis et al. 2023; keiding et al. 2024; westgate et al. 2024, 2025). 4.1.5. thorning structure the thorning structure is a three-way dip closure, elongated in a ssw–nne direction, located in central jylland (figs 1, 12e). the main influence on the structure is the thorning salt pillow, which affects the overlying triassic to cretaceous successions (fig. 12e). the salt pillow has recently been mapped and defined with the newly acquired data (bjerager et al. 2024). there are no deep wells in the structure, and the new seismic data acquired in 2023 (putnaite & malehmir 2024) are thus correlated to nearby wells to the south-west (nøvling-1 and vinding-1) and to the north (kvols-1) using legacy data. these wells are located some 30–40 km away from the thorning structure (figs 4, 5). the primary reservoir-seal pair for the thorning structure consists of the gassum and fjerritslev fms (hjelm et al. 2022; bjerager et al. 2024). a geological cross-section through the thorning structure is shown in figure 12e. deeper sandstones of the triassic skagerrak and tønder fms and shallower reservoirs, possibly the uppermost jurassic to lower cretaceous frederikshavn fm, known from wells to the north (e.g. kvols-1, hobro1, gassum-1; fig. 4), may form secondary reservoirs in the structure (bjerager et al. 2024). the gassum fm is c. 100 m thick on average, and the shallowest point of the reservoir within the structure is 1520 m bmsl. the lowest closing contour for the reservoir is at 1950 m bmsl, outlining an area of 235 km2 (bjerager et al. 2024). reservoir properties of the gassum fm sandstones are inferred to be good from nearby wells surrounding the thorning structure and from seismic interpretation estimated in three scenarios, showing an average effective reservoir porosity of c. 26–29% and an average permeability of c. 1000–1500 md (bjerager et al. 2024). the thorning structure overlies a salt pillow that began to form during the triassic and developed through the jurassic to early cretaceous, when continuous burial by thick sedimentary successions and tectonic activity probably triggered salt migration (bjerager et al. 2024). the salt pillow overlies a slightly northward-dipping pre-zechstein surface. a few extensional faults are observed in the triassic to cretaceous successions at the northern rim of the thorning structure. these faults were probably formed during the triassic and jurassic as a result of salt movements. compared with the earlier interpretation based on only a few seismic lines (fig. 5), the updated interpretation using the newly acquired high-resolution seismic data shows a more elongate structure (bjerager et al. 2024; putnaite et al. 2025). specifically, the structure is narrower in the e–w direction and longer in the n–s direction than originally mapped (hjelm et al. 2022; fig. 1). estimates of the size of the thorning structure are similar to those derived in previous studies (hjelm et al. 2022; bjerager et al. 2024). near-surface structures, including faults and palaeo valleys, are also revealed by the new seismic data (putnaite et al. 2025). a significant graben structure, with normal faults and a width of 5 km (south of the profile in fig. 12e), is revealed in the new seismic data (lines p4 and p5) at the sw flank of the thorning structure (see bjerager et al. 2024). it probably formed due to salt withdrawal and collapse, resulting in a marked thickening of the fjerritslev fm in the fault zone. large throws at the top gassum to top falster horizons indicate faulting during deposition of the vinding fm, which shows a clear increase in thickness near to the ne-dipping main boundary fault of the graben structure. this was most likely the result of salt withdrawal beneath the graben. a second phase of faulting and salt migration took place during deposition of the fjerritslev fm. the top fjerritslev to base chalk horizons show only minor faulting, which probably developed during the growth of the salt pillow, partly triggered by the mid-cimmerian tectonic phase. thinning of the fjerritslev fm over the top of the structure may indicate elevation of the structure and associated erosion during the late early jurassic – early cretaceous. onlap of the lower cretaceous vedsted fm onto the top of the fjerritslev fm marks a major hiatus at the mcu. the fault zone along the sw flank of the thorning structure is close to the deepest closing contour, that is, the spill point for the top gassum fm of the structure. from the point of view of leakage risk, this is therefore less critical than if it had occurred at the top of the structure. the profiles in the new seismic survey are rather sparse; new and https://doi.org/10.34194/299dt488 gregersen et al. 2025: geus bulletin 60. 8385. https://doi.org/10.34194/299dt488 21 of 29 geusbulletin.org more tightly spaced seismic data are thus needed to further map and evaluate faults and determine their significance for seal integrity. 4.1.6. jammerbugt structure the jammerbugt structure (fig. 12f) is an elongated structure with a se–nw direction located in the fjerritslev trough near the shore of the jammerbugt bay, west of northern jylland (fig. 1). the structure outlines an undrilled, faulted, three-way closure. prior to the new seismic survey in 2023, the structure was only covered by a few seismic lines of variable quality. the new acquisition added 1450 km of seismic lines to the database, covering the structure and the adjacent area in a systematic and dense grid (figs 5, 9; table 1). three of the lines tie to the nearby j-1 well in the adjacent lisa structure (fig. 9). the new data permit a better delineation of the structure and an analysis of its geological evolution. extensional faults confine the north-western and north-eastern flanks of the jammerbugt structure, while the opposite flanks are defined by the stratigraphic plunge (fig. 12f; fyhn et al. 2024). seismic data are essentially lacking in the south-easternmost, landward part of the structure due to the low water depths there. the jammerbugt structure developed together with the fjerritslev trough that formed in response to mesozoic extension and down-throw across the fjerritslev fault. similarly, the faults in the jammerbugt structure developed in response to triassic–cretaceous pulses of deep-seated extension (fyhn et al. 2024). in addition, detachment faults in zechstein salt caused roll-over folding in the overlying mesozoic section, which influenced the structural architecture giving rise to an anticlinal element in the closure geometry. apart from the confining faults, the stratigraphy within the jammerbugt structure, including reservoirs and seals, is offset by faults rooted in the salt of the oddesund fm (indicated in fig. 12f). growth of a salt pillow within the oddesund fm under the north-western part of the structure contributed to the closure relief in the overlying section. similarly, doming associated with late cretaceous to palaeogene structural inversion of the fjerritslev trough contributed slightly to the closure relief. correlation to wells drilled in the vicinity suggests the presence of two reservoir levels within the structural closure: (1) the gassum fm and (2) the haldager sand fm. depth-converted seismic mapping of the jammerbugt structure places the shallowest point of the gassum fm at around 1620 m and the top of the haldager sand fm at 1160 m. the area of the closures within the gassum fm and haldager sand fm are mapped to be 119 km2 and 142 km2, respectively (fyhn et al. 2024). by comparison to nearby wells, the gassum fm is estimated to have a thickness of around 200 m, whereas the haldager sand fm is only estimated to have a thickness of around 20 m over the structure. the two reservoir intervals are interpreted to be overlain by thick mudstone successions. seismic mapping and comparison to nearby wells suggest that the mudstone-dominated lower jurassic fjerritslev fm, which overlies the gassum fm, has a thickness of a few hundred metres within the structure. the fjerritslev fm forms the primary seal for the gassum fm. the haldager sand fm is interpreted to be overlain by upper jurassic børglum fm claystones that form a sealing unit. information from nearby wells intersecting this finegrained unit together with seismic mapping suggests a thickness of the børglum fm of c. 100 m in the jammerbugt structure. a geological risk defined at this stage is associated with faulting of the reservoir and seal intervals, where some faults appear to terminate close to the seabed. these faults introduce a risk of reservoir leakage and compartmentalisation and hence the potential risk for leakage along fault planes needs further investigation. other geological uncertainties include the local reservoir quality of the gassum fm within the structure and the risk that hydrocarbons, formed in the fjerritslev trough, have charged the undrilled structure. 4.1.7. lisa structure the lisa structure is located in the fjerritslev trough, offshore northern jylland (fig. 1). it is a four-way dip closure elongated in a se–nw direction; the structure is located above a salt pillow belonging to the oddesund fm (fig. 12g). the structure was drilled by the j-1 well and is covered by a sparse grid of seismic data of variable quality acquired by both industry and academia. the most recently acquired data over the structure was acquired in 2023 as part of this study (fig. 5). the j-1 well terminated in rhaetian deposits after drilling 1952 m of upper cretaceous, jurassic, and uppermost triassic strata. the lisa salt pillow is overlain by two reservoir-seal pairs, the gassum fm – fjerritslev fm and the haldager sand fm – børglum fm, where four-way dip closures are present at both stratigraphic levels. the structures in the sequence overlying the salt developed in response to a combination of differential salt motion and late cretaceous to palaeogene structural inversion of the fjerritslev trough. the rhaetian to hettangian gassum fm forms the primary reservoir with a thickness of 199 m in the j-1 well. based on calculations using the wireline logs from this well, the reservoir is estimated to have a net-togross of 0.45, an average effective reservoir porosity of 20% and an average permeability of 251 md. the gassum fm is capped by 623 m of mudstone-dominated strata belonging to the lower jurassic fjerritslev fm of which the lower 120 m section, dominated by shale, is https://doi.org/10.34194/299dt488 gregersen et al. 2025: geus bulletin 60. 8385. https://doi.org/10.34194/299dt488 22 of 29 geusbulletin.org considered an excellent main seal for storage within the gassum fm. despite an excellent average effective reservoir porosity of 25% and permeability of 1112 md the shallower middle jurassic haldager sand fm is considered a secondary reservoir due to its modest thickness and net-to-gross of 19 m and 0.24, respectively. the haldager sand fm reservoir is overlain by 101 m of mudstones belonging to the børglum fm, which are considered to have excellent seal potential. both reservoirs and seals are intersected by faults, some of which appear to terminate close to the seabed. over part of the closure, faults are densely spaced, located at intervals of a few kilometres to a few hundred metres apart. these faults typically offset the geological layers by a few tens of metres but occasionally more. the primary geological risks for efficient and lasting co2 storage as identified at this stage are therefore associated with the presence of these minor faults offsetting both reservoirs and overlying seals. the faults introduce the risk of reservoir compartmentalisation and a mechanical weakening of the seal, which needs to be investigated with further data acquisition and analysis. 4.1.8. inez structure the inez structure is situated c. 50 km offshore in the northern danish north sea (fig.1). it is a four-way dip structure outlining one of the larger danish geological structures with a promising storage potential. the inez area is covered by a sparse 2d seismic grid of variable quality, and the structure was drilled in 1978 by the inez-1 well. the inez structure (fig. 12h) is a turtleback structure formed in response to zechstein salt migration towards six large salt structures in the vicinity of the structure, where salt movement occurred predominantly during the middle triassic to early cretaceous. the associated differential subsidence resulted in laterally migrating rim synclines around the inez structure generating a four-way dip closure at three reservoir levels: (1) top skagerrak fm, (2) top gassum fm and (3) top haldager sand fm (fig. 12h). the gassum and haldager sand fms are intersected by the inez-1 well and include sandstones with average effective porosities of 20.3% and 26.0%, respectively. estimated average permeabilities of the gassum and haldager sand fms are 442 md and 871 md, respectively. the rhaetian to hettangian gassum fm, which formed in a near-shore environment, is interpreted to have a thickness of 148 m in the inez-1 well. it has a net-to-gross ratio of around 0.59. the middle jurassic haldager sand fm only measures 9 m in thickness, has a net-to-gross ratio around 0.32 and is therefore secondary to the gassum fm. the gassum fm is overlain by thick claystone intervals of the fjerritslev fm, which are observed to be 127 m thick in the inez-1 well. hence, the fjerritslev fm is considered the primary seal for the gassum fm reservoir. similarly, the jurassic sandstone interval (haldager sand fm) is overlain by thick upper jurassic claystones of the børglum fm, which form a sealing unit. in the inez-1 well, the youngest gassum fm is lowermost jurassic in age. the section downwards becomes increasingly arkosic in nature interpreted as an upper triassic proximal facies belonging to the gassum or vinding fms in which the well terminates. the lower to lowermost upper triassic skagerrak fm is correlated seismically from the felicia-1a well with the deeper part of the inez structure. the thick and sand-rich skagerrak fm comprises another secondary reservoir in this structure. the shallowest point of the skagerrak fm within the structure is estimated to be of a depth of about 2500 m, which lies some 175 m above the lowest closing contour. the unit is not intersected in the inez-1 well as the well terminated in the uppermost triassic succession, but good reservoir properties are anticipated from analogues within danish wells elsewhere that encounter the skagerrak fm. the skagerrak fm is interpreted to be overlain by fine-grained sediments and evaporites of the oddesund fm that probably provide a tight sealing unit for the reservoir. a geological risk defined at this stage is associated with densely spaced faulting within the gassum fm; these faults typically offset the reservoir by a few tens of metres introducing a risk of reservoir compartmentalisation. the faults also continue upwards into the overlying fjerritslev fm seal succession. therefore, although the fjerritslev fm seal thickness fully complies with the recommendations for co2 storage, the potential risk for leakage along fault planes needs further investigation. 4.2. new geological results and implications the study of the eight structures in the ccs2022–2024 project has resulted in an improved geological understanding of the structures and basins, their tectonostratigraphic evolution and the regional correlation between the structures. it has also revealed several novel questions to be addressed in future work. the newly acquired seismic data significantly improve the existing database and allow for a revised and much more detailed interpretation with a better physical definition of the structures, as well as new depth-structure maps of key horizons and thickness maps. the new data and interpretation also provide valuable additional information regarding the composition of reservoir and seal successions, faults and biostratigraphy, all of which are important for future work on the suitability of these structures for the purpose of ccs. https://doi.org/10.34194/299dt488 gregersen et al. 2025: geus bulletin 60. 8385. https://doi.org/10.34194/299dt488 23 of 29 geusbulletin.org in all structures, the primary reservoir formations (storage formations) show a significant and relatively stable thickness, mostly between 100 and 200 m for the gassum fm and 200 to 250 m for the bunter sandstone fm (table 2). good reservoir properties are demonstrated based on interpretation of well data in or near the structures. there are no wells drilled in the havnsø, thorning and jammerbugt structures; the reservoir models of these structures are based on interpretation of data from the wells closest to the structures and interpretation from seismic tie-lines correlated from these wells. we interpreted seismic facies successions with subtle inclined reflections and troughs, suggestive of sandstone-dominated lowstand successions. this was particularly the case for the havnsø structure, based on similar features and seismic correlation from the stenlille area with wells and 3d seismic data (gregersen et al. 2023b). in all structures, the primary seal formation is based on wells in or near the structures and mapped from seismic data that demonstrate significant thicknesses. the fjerritslev fm is more than 250 m thick in most structures, except for the inez structure, which has a thickness of 127 m. in the rødby structure, the ørslev fm seal succession has a thickness of 175 m (table 2). the structures mainly formed through episodic salt tectonism triggered by deep-seated faulting. further investigations of structures, reservoirs and seals are recommended to increase the understanding of their co2 storage potential. previously unobserved faults, most significantly in the jammerbugt, gassum, rødby and thorning structures, have been discovered in the new seismic data. however, faults are recognised in both new and vintage data in all structures. therefore, all investigated structures require additional seismic and well data to perform an assessment of risks associated with potential co2 storage, in particular regarding faults, but also regarding old wells where they penetrate seal and reservoir successions. interpretation using regional seismic correlation between wells and analysis of well data during this project have revealed a need for updating the lithostratigraphy in the triassic succession of the danish basin. different formation names (e.g. the skagerrak fm and the bunter sandstone fm) are in some cases used for equivalent stratigraphic intervals. in addition, regional lithofacies distribution (e.g. of triassic reservoir sandstones) is not well understood, while existing lithostratigraphy is based on wells but not integrated with regional seismic correlations. therefore, a revision of the danish triassic stratigraphy should be considered. 4.3. storage capacity estimation using the new high-quality seismic data, the work in the ccs2022–2024 project has resulted in an improvement in the physical definition of the eight structures, their outline areas and grvs (fig. 12). the new geological knowledge of the structures has increased the understanding of the reservoir and seal successions. the new data and interpretations have also significantly improved the understanding of the geometry and distribution of faults. this improved knowledge is important for further ccs evaluation. in this study, updated well analysis and new depth conversions have also led to a greatly improved characterisation of key reservoirs and the geometry of the structures. evaluation and maturation of a co2 storage site include several steps. a maturation phase includes a static calculation of the theoretical storage capacity  – primarily table 2 area, depth and formation thicknesses of each of the structures in the ccs2022–2024 project. structure area top primary reservoir fm primary seal fm (km2 at top (m bmsl to top formation thickness formation thickness primary reservoir) primary reservoir) (m) (m) stenlille 5.4 km2 1449 m gassum fm 150 m fjerritslev fm 275 m havnsø 70 km2 1550 m gassum fm 150 m fjerritslev fm 275 m rødby 117 km2 1100 m bunter sandstone fm 230 m ørslev fm 175 m gassum 280 km2 1375 m gassum fm 180 m fjerritslev fm 325 m thorning 235 km2 1520 m gassum fm 94 m fjerritslev fm 265 m jammerbugt 119 km2 1620 m gassum fm 200 m fjerritslev fm 525 m lisa not available 1623 m gassum fm 199 m fjerritslev fm 623 m inez not available 1592 m gassum fm 148 m fjerritslev fm 127 m ‘area’ refers to the area within the lowest closing contour for each respective structure, based on a depth-converted structure map of the top of the reservoir. ‘top’ refers to the shallowest point of the primary reservoir within the structure. the table shows simplified single values to give an overview for comparison between the structures. the values are mostly averages from a ranges of scenarios, based on wells and mapping of the individual structures, which are more thoroughly described in the interpretation reports of the project (gregersen et al. 2023a, 2023b; abramovitz et al. 2024; bjerager et al. 2024; fyhn et al. 2024; keiding et al. 2024). note that the formation thicknesses from maps change across the structures and become thicker mainly from the top and towards the flanks of the structures (fig. 12) as described in the interpretation reports of the project (see references above). depth and thicknesses for each structure are from wells and maps, although the values for the lisa and inez structures were derived solely from the j-1 and inez-1 wells, respectively. https://doi.org/10.34194/299dt488 gregersen et al. 2025: geus bulletin 60. 8385. https://doi.org/10.34194/299dt488 24 of 29 geusbulletin.org based on grv, reservoir sand thickness and average effective porosity, as well as density of the co2 (section 3.6) but excluding permeability. the current maturation phase does not include dynamic capacity estimates of the potential co2 structures but focuses on identifying and assessing the extent and quality of the reservoir aquifers. furthermore, no attempts are made to address seal capacity (i.e. co2 entry pressures), fault leakage, fault reactivation, solubility of co2 in brine, and co2 mineralisation reactions. by using the static approach (described in section 3.6), all the structures of this project have been re-evaluated since hjelm et al. (2022). the gassum fm reservoirs have been the primary sandstone target with the bunter sandstone fm reservoir in the rødby structure being the exception. sandstone reservoirs in the frederikshavn, haldager sand and skagerrak fms may also possess considerable reservoir potential, but the storage potential of these units has not been evaluated. combined, these units may provide a significant upside to the storage capacity. in the storage capacity estimation, the grv is corrected with the n/g ratio to achieve a more realistic reservoir sand volume. extended and more detailed analysis of the log data show that the gassum fm in some wells has a number of discrete sandstone intervals, mainly in the lower part of the formation (e.g. the havnsø and thorning structures). however, these separate layers are not taken into account in these estimations. to evaluate the uncertainty on the input parameters, minimum and maximum cases were also calculated by assigning a minimum, mode and maximum uncertainty range, mode being the data value that occurs most often in the data set. it is assumed that the assigned distribution of all the input parameters follow a pert distribution defined by the minimum, mode and maximum values. the pert distribution is believed to give suitable representation for naturally occurring events following the subjective input estimates (clark 1962). the variation in grv, amongst others, was inferred to cover uncertainty in interpretations, seismic well ties, mapping and depth conversion normally by defining the minimum and maximum of the distribution based on surrounding wells. some variation of n/g and porosity are expected due to thickness and lateral variation of the lithologies owing to differences in facies distribution, depositional environment, diagenesis and poor quality of the well logs. storage efficiency is heavily influenced by local geological subsurface factors, and an analogue storage efficiency database is not available for the danish onshore area; accurate storage efficiency factor ranges are thus lacking at this early stage of maturation. a fixed factor value of 0.1 was used, assuming that the reservoir has reasonable reservoir characteristics, and that uncertainty is caused by the identification of faults on or near the apex or top point of the structures penetrating both the seal and the reservoir. the results of monte carlo simulations for each of the structures, including the mean mass of co2 in megatons (mt) that can be stored, are shown in table 3. together with the 90, 50 and 10% percentiles (p90, p50 and p10), the range corresponds to the chance for a given storage volume scenario to exceed the given storage capacity value. mean values of the resultant outcome distribution is considered the ‘best’ single value representation for the entire distribution. without addressing the influence of the faults located on or near the apex or top point of several of the structures (e.g. the gassum and rødby structures), the mean unrisked static storage capacities for the investigated structures range between 8 and 498 mt co2 for the gassum fm, and 107 mt co2 for the bunter sandstone fm (table 3). due to the variability of the underlying factors, the estimated storage capacities have a significant range. the results are mainly linked to the use and uncertainty table 3 unrisked theoretical storage capacity for structures investigated in the ccs2022–2024 project. structure stratigraphic level area (km2) static co2 capacity mean (mt co2) p90 (mt co2) p50 (mt co2) p10 (mt co2) stenlille gassum fm 5.4 8 1.7–4.7 2.5–7.7 3.2–11.8 havnsø gassum fm 70 50 22.1–41.3 33.5 – 62.8 48.8–90.4 rødby bunter sandstone fm 117 107 69 104 149 gassum gassum fm 280 498 325 486 689 thorning gassum fm 235 125 81 122  174  jammerbugt gassum fm 119 199 122 191 289 the primary reservoir formation is the gassum fm except for the rødby structure, where the bunter sandstone fm is the primary reservoir formation. for the stenlille structure, a fixed storage efficiency factor of 0.4 was used, whereas for all other structures, a fixed efficiency factor of 0.1 was used. the capacity ranges for the stenlille and havnsø structures cover different scenarios, and the mean values cover an average of scenarios: 6–10 mt co2 and 35–65 mt co2 for the stenlille and havnsø structures, respectively. estimations for the lisa and inez structures are in progress and were not available at the time of publication. p90, p50 and p10: 90, 50 and 10% percentiles, respectively. see full estimations in the interpretation reports of the project (gregersen et al. 2023a, 2023b; abramovitz et al. 2024; bjerager et al. 2024; fyhn et al. 2024; keiding et al. 2024). https://doi.org/10.34194/299dt488 gregersen et al. 2025: geus bulletin 60. 8385. https://doi.org/10.34194/299dt488 25 of 29 geusbulletin.org of the storage efficiency factor, where the factor is 0.1 (minimum of 0.05 to a maximum of 0.2) compared to the previously used 0.4, thus reducing the overall co2 storage capacity compared to previous studies (see hjelm et al. 2022). the simple method used to obtain the estimated storage capacities in table 3 is a means to benchmark and compare the structures during the initial screening across denmark. it should be updated when more data are gathered. it is also recommended to apply more complex and detailed dynamic methods. using a simple static co2 storage capacity calculation is, however, beneficial for several reasons. it only requires a few input parameters, making it easier to obtain first estimates on capacity. it is cost-effective, less expensive and less time-consuming than dynamic methods, which require more data and computational power. it is important to note that static methods do not account for fluid flow and more complex geological factors and facies configurations. furthermore, static calculations do not account for all the dynamic processes that can affect co2 injection and storage over time. 5. conclusions and perspectives the ccs2022–2024 project has matured eight selected geological structures (the gassum, havnsø, inez, jammerbugt, lisa, rødby, stenlille and thorning structures) that were identified as suitable targets for potential storage of co2, facilitated by new seismic acquisition and hence enhanced interpretation. the seismic acquisition of the project carried out in 2022 and 2023 is one of the most comprehensive public scientific acquisition campaigns ever conducted in denmark, and it has significantly increased the data coverage with high-quality seismic data over the structures investigated. within the project, a total of 2093 km of new 2d seismic profiles of good quality were acquired, where approximately 643 and 1450 km were acquired onshore and offshore, respectively. the data sets were initially processed and subsequently reprocessed to enhance more details. in addition, selected pre-existing legacy seismic data sets were reprocessed. the new extended database was used to update the geological and seismic interpretation and structural mapping of the structures, providing both a better delineation of the investigated structures, including major faults, and an improved understanding of the geological and tectonostratigraphic evolution. interpretation of seismic facies reflecting sandstone-dominated successions away from wells support the interpretation that storage formations are regional in extent. all new data along with the acquisition, processing and interpretation reports and a list of related publications from the ccs2022–2024 project are available on the geus website (https://geus.dk/ccsdata). key results of the project and continued work are summarised within this bulletin. unrisked theoretical static storage capacities of the structures were estimated based on the new data and interpretations performed in the project and are provided with ranges of uncertainty. the assessment of the static estimated storage capacities provides a good insight into the relative capacities of the structures. however, they do not address complex variations in the reservoir structure and properties. hence, this highlights the need for further work, for example to identify volumes of the reservoir with high quality, connectivity and absence of cross-cutting faults. after this maturation phase, the reservoir units and the impact of identified faults must be assessed in more detail by dynamic 3d reservoir simulation models to ensure optimal development, injection, and filling of the storage reservoirs of the structures. in addition, the storage capacity should be determined with a lower uncertainty. the results of the project also lay the foundation for developing new projects including seismic surveys for further evaluation of the structures for ccs and other purposes. new and more densely spaced data such as 3d seismic surveys are needed for the further maturation steps to fill in the data gaps and to address 3d complexities, such as to further resolve and describe the structural traps, reservoirs, seals, spill points, faults, storage capacity and to address potential risks. in addition, monitoring of for example seismicity, groundwater, wells, faults, elevation and other investigations are very important prior to, during and after co2 storage operations. the project has revealed that not only siteand structure-specific work is required, but also that more regional geological studies should be carried out to revise the regional understanding of the depositional systems, including reservoir distribution, seal characterisation and tectonostratigraphic evolution across basins and large structures of the danish basin. in addition, the stratigraphy and geological evolution from the danish basin to adjacent regions should be investigated, southwards across the ringkøbing–fyn high, northwards into the sorgenfrei–tornquist zone and skagerrak–kattegat platform, eastwards into the øresund basin and westwards into the norwegian–danish basin. the new data and related research have resulted in numerous scientific publications and the completion of several phd and master theses, which benefited from using the new data and participating in the field work. this is also an outstanding societal contribution to the education of a new generation of geoscientists working to tackle climate issues, and to the green energy transition. https://doi.org/10.34194/299dt488 https://geus.dk/ccsdata gregersen et al. 2025: geus bulletin 60. 8385. https://doi.org/10.34194/299dt488 26 of 29 geusbulletin.org acknowledgements the enthusiastic and fruitful cooperation between the authors, and with the external parties of geopartner geofizyka (poland) and cowi (denmark) in connection with the seismic data acquisition is much acknowledged. we are grateful for the hard work of the field crews, including students from uppsala university (sweden), university of copenhagen (denmark), aarhus university (denmark) and unilasalle (france). we also thank per trinhammer for cooperation regarding the marine acquisition at havnsø, tomi a. jusri and rasmus rasmussen for qc of the seismic data processing, and lasse m. rasmussen for data support. we appreciate the careful figure drafting by carsten e. thuesen. annette ryge is thanked for providing biostratigraphic preparations. we acknowledge the constructive and useful comments and suggestions by the two reviewers. this study is part of the ccs2022–2024 project for maturation of selected structures to potential co2 storage sites. additional information funding statement the funding to geus for the maturation of the eight selected, potential co2 storage sites in denmark (the ccs2022–2024 project) is provided by the ‘funding of higher education in denmark’ (forskningsreserven 2022). the funding made the contracted seismic acquisition and the subsequent work including this publication possible. author contributions ug: conceptualisation; writingoriginal draft; main 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http://www.peacesoftware.de/einigewerte/co2_e.html http://www.peacesoftware.de/einigewerte/co2_e.html https://doi.org/10.1029/2022ea002464 https://doi.org/10.1190/geo2023-0503.1 seismic investigations of eight geological structures for potential storage of co2 in denmark: an introduction 1. introduction 2. geological setting 3. database and methods 3.1. database 3.2. onshore seismic acquisition 3.3. offshore seismic acquisition 3.3.1. offshore: havnsø-nekselø seismic acquisition 3.3.2. offshore: jammerbugt seismic acquisition 3.4. seismic interpretation and mapping 3.5. investigation of reservoir and seal 3.6. storage capacity 4. results of seismic interpretation and storage capacity estimation 4.1. stratigraphy, geometry and geological development of individual structures 4.1.1. stenlille structure 4.1.2. havnsø structure 4.1.3. rødby structure 4.1.4. gassum structure 4.1.5. thorning structure 4.1.6. jammerbugt structure 4.1.7. lisa structure 4.1.8. inez structure 4.2. new geological results and implications 4.3. storage capacity estimation 5. conclusions and perspectives acknowledgements additional information funding statement author contributions references figures fig. 1 map of the danish structures with potential for geological storage of co2 from the geus-led ccs2022–2024 project. the dark green structures (stenlille, havnsø, rødby, gassum, thorning, jammerbugt, lisa and inez structures) are mapped in the project. outlines of the lisa and inez structures are modified from hjelm et al. (2022). the dark green shading shows the extent of the deepest mapped closure of the top gassum fm surface, except in the rødby structure, where the deepest closure of the top bunter sandstone fm is delineated. the light green shading shows outlines of other structures, which may have potential for geological storage of co2. fig. 2 top pre-zechstein map of the main structural elements onshore and offshore denmark, including highs, basins and main faults. the elements include the norwegian–danish basin, the eastern part of which is named the danish basin, the sorgenfrei–tornquist zone, the skagerrak–kattegat platform, the ringkøbing–fyn high and the northern part of the north german basin. modified from vejbæk (1997). fig. 3 schematic stratigraphic diagram oriented from north to south including wells in the sorgenfrei–tornquist zone (stz), the eastern part of the danish basin (db), ringkøbing–fyn high (rfh) and the north german basin (ngb), compiled in this study. the well sections are from north to south located geographically in the kattegat sea (terne-1), and on the islands of sjælland (stenlille-19, slagelse-1), fyn (ullerslev-1), falster (ørslev-1) and lolland (søllested-1, rødby-2) shown by red vertical lines (arrows indicate where well termination, total depth, is located below the figure). well locations are shown in fig. 4. the lithostratigraphy is based on available well sections (nielsen & japsen 1991 and references therein) and new research from this project using wells and seismic data. similar groups and formations occur in jylland, and the figure is used as a reference for the current danish lithostratigraphy (incl. bertelsen 1978, 1980; michelsen et al. 2003; nielsen 2003). the upper permian-triassic lithostratigraphy of northern germany is also shown. modified from abramovitz et al. (2024). fig. 4 locations of danish wells (red circles) and outlines of regional structural elements, including structural highs, basins, fault zones and faults. modified from nielsen (2003). fig. 5 locations of the new seismic lines (red) acquired in 2022 and 2023, and outlines (green) of the structures that were mapped as part of this project (ccs2022–2024). older seismic lines are shown in blue. fig. 6 examples of vintage and newly acquired 2d seismic data shown with vertical scale in two-way travel (twt) time in milliseconds (ms). example from onshore seismic data across the havnsø structure with (a) vintage seismic data (ref: ssl6267-r12) and (b) new seismic data acquired and processed by uppsala university (ref: geus22-hvn-p7). example from offshore seismic data across the jammerbugt structure with (c) vintage seismic data (ref: wgc64a-39970) and (d) newly acquired seismic data by the federal institute for geosciences and natural resources (bgr) and aarhus university, and that was subsequently reprocessed by realtimeseismic (rts; ref: geus23-jb-15 from the reprocessed survey geus2023-jammerbugt-re2023). fig. 7 photographs from the seismic acquisition with the two vibroseis trucks, landstreamer and geophones. (a) operational setup of field equipment. (b) mems sensors mounted at 2 m intervals on a landstreamer towed behind the rear vibroseis truck. (c) two vibroseis trucks are operated with synchronised vibrations. (d) wireless geophones are deployed every 10 m along the profile. reproduced from malehmir & westgate (2023). fig. 8 overview of the havnsø seismic acquisition from havnsø to nekselø. (a) location map showing the onshore seismic sources (vibroseis) and the marine receivers (obs: ocean bottom seismometers, and streamer cable) used along the northern portion of seismic line geus22-hvn-p1. (b) schematic cross-section of the marine receivers. (c) the streamer winch on the beach at havnsø with 600 m of streamer cable deployed at the seafloor towards nekselø. (photograph by per trinhammer 2022) (d) four of the retrieved obs on board a local fishing vessel that was used for the marine operation. (photograph by egon nørmark 2022). fig. 9 topographic map with the location of the acquired seismic lines of the geus23-jb survey (red lines). yellow circles indicate the positions of wells, the black line marks the track of the acquisition vessel jákup sverri, blue lines indicate the 10 m depth contour. fig. 10 photographs from the jammerbugt seismic acquisition. (a) the faroese research vessel jákup sverri about to leave hirtshals harbour (northern jylland) for the seismic acquisition. (b) deployment of the streamer tail buoy. in the foreground, one of the two winches with the solid-state streamer cable (sercel sentinel ssrd; yellow cable). photographs: thomas funck. fig. 11 conceptual profile (a–a’) across a potential structure. the uncertainty in mapping the structure results in the hypothetical minimum (min.) and maximum (max.) scenarios that are very different from the most likely mapped scenario. variance in area and in gross thickness (t) will affect the gross rock volume (grv) of the structure. the uncertainty is addressed by applying uncertainty on the resulting grv and other parameters and by conducting simple monte carlo simulation to calculate 90, 50 and 10% percentiles (e.g. burruss et al. 2009; heidug et al. 2013). fig. 12 simplified geological cross sections through the eight structures investigated in this study. the approximate locations of the largest faults are shown. the primary reservoir formation is the gassum fm for all structures except for the rødby structure (c), where it the primary reservoir formation is the bunter sandstone fm. (a) stenlille structure. (b) havnsø structure. (c) rødby structure. (d) gassum structure. (e) thorning structure. (f) jammerbugt structure. (g) lisa structure. (h) inez structure. note that the sections have different vertical and horizontal scales. tables table 1 seismic data (2d) acquired 2022 and 2023 during the ccs2022–2024 project and pre-exisiting data. table 2 area, depth and formation thicknesses of each of the structures in the ccs2022–2024 project. table 3 unrisked theoretical storage capacity for structures investigated in the ccs2022–2024 project. geological survey of denmark and greenland bulletin 11, 145-161 145 the nordre strømfjord shear zone and the arfersiorfik quartz diorite in arfersiorfik, the nagssugtoqidian orogen, west greenland kai sørensen, john a. korstgård, william e. glassley and bo møller stensgaard the nordre strømfjord shear zone in the fjord arfersiorfik, central west greenland, consists of alternating panels of supracrustal rocks and orthogneisses which together form a vertical zone up to 7 km wide with sinistral transcurrent, ductile deformation, which occurred under middle amphibolite facies conditions. the pelitic and metavolcanic schists and paragneisses are all highly deformed, while the orthogneisses appear more variably deformed, with increasing deformation evident towards the supracrustal units. the c. 1.92 ga arfersiorfik quartz diorite is traceable for a distance of at least 35 km from the inland ice towards the west-south-west. towards its northern contact with an intensely deformed schist unit it shows a similar pattern of increasing strain, which is accompanied by chemical and mineralogical changes. the metasomatic changes associated with the shear zone deformation are superimposed on a wide range of original chemical compositions, which reflect magmatic olivine and/ or pyroxene as well as hornblende fractionation trends. the chemistry of the arfersiorfik quartz diorite suite as a whole is comparable to that of phanerozoic plutonic and volcanic rocks of calc-alkaline affinity. keywords: aeromagnetic data, arfersiorfik quartz diorite, deformation, geochemistry, nagssugtoqidian, nordre strømfjord shear zone, west greenland ________________________________________________________________________________________________________________________________________________________________ k.s. & b.m.s., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ks@geus.dk j.a.k., department of earth sciences, university of aarhus, dk-8000 århus c, denmark. w.e.g., lawrence livermore national laboratory, l-646, livermore, ca 94550, california, usa. the c. 1850 ma nagssugtoqidian orogen in central west greenland, originally defined by ramberg (1949), is built up of both palaeoproterozoic and reworked archaean rocks, and contains several prominent shear zones of nagssugtoqidian age (bak et al. 1975a; connelly et al. 2000; van gool et al. 2002). these are roughly ene–wsw-trending, approximately linear zones of intensely deformed rocks with well-developed planar and linear tectonic fabrics. the shear zones may be up to 20 km wide and have a strike length of more than 150 km, the approximate width of the ice-free coast of west greenland. one of the most prominent shear zones, the nordre strømfjord shear zone (fig. 1), was described in detail by bak et al. (1975b) and sørensen (1983), and is a main subject of this paper. the nordre strømfjord shear zone was recognised by previous authors as a sinistral shear zone mainly on the basis of deflection of fabrics and lithologies. a general observation was that open structures outside the zone become closed and brought into parallelism within the zone. this is apparent both in the field and especially so on regional geological maps and aerial photographs. a systematic change in the orientation of planar fabrics across the shear zone indicated that the zone is wedge-shaped in vertical profile, suggesting that the zone narrows upwards, © geus, 2006. geological survey of denmark and greenland bulletin 11, 145–161. available at: www.geus.dk/publications/bull 146 disko bugt 69° –0.844 –0.230 –0.147 –0.108 –0.082 –0.064 –0.049 –0.037 –0.027 –0.019 –0.011 –0.004 0.004 0.013 0.024 0.039 0.059 0.087 0.128 0.210 1.736 attu arfe rs iorfik 25 km inland ice b [nt/m] b aasiaat attu nordre strømfjord shear zone nordre strømfjord arfersiorfik fig. 1b fig. 2 agto map sheet 54° 68° a 50 km inland ice 53°53° 51°51° 50°50°52°52° 51° 50°52° 67°30'67°30' 53° 67°30' 51° 68°68°68° arfe rs iorfik 68° arfe rs iorfikattuattu 500 km greenland surficial deposits quaternary basalt palaeogene metasedimentary rocks (may include archaean components) nagssugtoqidian orogen sisimiut charnockite arfersiorfik quartz diorite orthogneiss (archaean, reworked) archaean metasedimentary rocks (may include palaeoproterozoic components) amphibolite anorthosite and ultrabasic rocks fig. 1. a: simplified geological map of the nagssugtoqidian orogen. index map of greenland shows the location of fig. 1a. black frames, locations of the agto map sheet (olesen 1984) and the inner arfersiorfik region investigated during field work in 2002 (fig. 2). b: aeromagnetic map of the nordre strømfjord shear zone from the coast to the western margin of the inland ice (shown as the vertical gradient of the total magnetic field intensity, nt/ m). anticlockwise rotation of lithologies into the shear zone is shown by curved stippled lines. the established and inferred boundaries of the high-strain part of the shear zone (full and stippled lines) are positioned where these rotating lithologies attain orientations which are indistinguishable from those within the central part of the shear zone itself. the magnetic anomaly field was obtained by subtracting the regional aeromagnetic data from the international geomagnetic reference field (rasmussen & van gool 2000). nominal flight altitude of the survey: 300 m above sea level, with gentle drape flying over areas with high relief. the survey was flown along n–s flight lines 500 m apart and along orthogonal tielines at 5 km intervals. nt = nanotesla (magnetic flux density). 147 and conversely widens deeper in the crust. this interpretation is supported by the observation that the deepest part of the shear zone, which is exposed in the west near the coast, is more than twice as wide as its more high-level, eastern end near the inland ice (bak et al. 1975a, b). the western part of the nordre strømfjord shear zone was mapped in detail within the agto map sheet (olesen 1984). in this paper we report an investigation of its eastern part, conducted in 2002 in the inner part of the fjord arfersiorfik. our investigations were in part intended to resolve conflicting interpretations of the nature of the shear zone. although all previous published work on the shear zone had reached similar conclusions regarding its magnitude and structural character, a different hypothesis was presented in a subsequent reconnaissance study by hanmer et al. (1997a, b), who re-interpreted the pre-existing data and purported that the shear zone was smaller than previously mapped and accommodated less lateral displacement than previously accepted. our premise was that detailed field work aimed at resolving the spatial relationships of structural elements, rather than relying on previously collected data, would expand our knowledge of the characteristics of the shear zone and would enable rigorous testing of the two previous hypotheses. the early work (bak et al. 1975a, b; sørensen 1983) showed the shear zone to be a large-scale synmetamorphic, sinistral transcurrent zone with a distinct wedge shape in its western, syn-granulite facies part. in that area, the shear zone is c. 15 km wide. it was also suggested that the shear zone deformation occurred simultaneously with large-scale metasomatism (sørensen & winter 1989). prior to the present study, the general outline and extent of the nordre strømfjord shear zone between the agto map sheet and the inland ice was drawn by bak et al. (1975b) and sørensen (1983), based on work by henderson (1969) and interpretation of aerial photographs. now, the extent of the zone of high strain between the coast and the inland ice can be demonstrated to correlate with anticlockwise rotation of lithologies into the shear zone, which is clearly visible on a new aeromagnetic data set (nielsen & rasmussen 2004). the age of the shear zone seems to be bracketed by the 1.85 ga age of regional metamorphism (hickman & glassley 1984; connelly et al. 2000) and postkinematic pegmatite ages of 1.76 ga (connelly et al. 2000; stendal et al. 2006, this volume). apart from the shear zone, the area we investigated in 2002 also encompasses a major part of the arfersiorfik quartz diorite (aqd), first observed by noe-nygaard & ramberg (1961) and later described by henderson (1969). the latter author visited the area in connection with reconnaissance work leading to the publication of the 1:500 000 scale geological map nussuaq – søndre strømfjord (escher 1971). the arfersiorfik quartz diorite was subsequently studied by kalsbeek et al. (1984, 1987) and kalsbeek (2001), who interpreted it as a subduction-related calc-alkaline intrusive suite and obtained an age of intrusion of 1.92 ga (see also the section on petrology). along its shores, the arfersiorfik fjord system offers excellent exposures of the shear zone and its surroundings. the authors spent three weeks in the region in the summer of 2002, conducting their investigations from a single camp with transport by inflatable dinghy, supplemented by half a day’s helicopter reconnaissance in the area between inner arfersiorfik and the inland ice. the area covered during the field work is shown in fig. 2. the main aims of the field work were to investigate the shear zone in the inner part of arfersiorfik and to investigate the relationships between the shear zone and the arfersiorfik quartz diorite; as can be seen from fig. 2, the shear zone constitutes the northern boundary of the aqd for a distance of at least 35 km between the inland ice and inner arfersiorfik. laboratory investigations of the approximately 110 samples collected during field work have not yet been finalised, and this account is therefore of preliminary nature. a transect through the shear zone the lithology and structure of the shear zone are illustrated by way of a schematic profile covering the area between akunnaaq and sarfaarsuk (fig. 3), a distance of c. 8 km, corresponding to a distance of c. 7 km as measured perpendicular to the shear zone trend. the shear zone is expressed by two principal characteristics: lithological units are all planar and vertically oriented, and the lithological variability per unit length across strike is much more pronounced than in the surrounding rocks north and south of the zone. in the field, these features are best displayed along the eastern shore of the small fjord orlerfik (fig. 2). lithologically, the shear zone can be described as a succession of supracrustal schist and gneiss units alternating with siliceous and intermediate orthogneisses of the aqd. the profile of fig. 3 is described in the following from south to north. further details about the aqd are described in a later section. to the south of the shear zone (within the area of investigation, fig. 2) the aqd is seen to rest with low-angle boundaries on siliceous gneisses, often with an intervening screen of schist. we agree with van gool et al. (1999) that this surface is of tectonic, probably thrust or ductile 148 thrust nature. farther to the south, in the area between the heads of nordre strømfjord and arfersiorfik, a thin (up to 500 m) sheet of aqd with tectonic contacts occurs with archaean siliceous gneisses both above and below the aqd sheet (the ussuit unit; manatschal et al. 1998; van gool et al. 1999). the geometry of this bounding surface is interpreted as outlining a series of folds. only the northern limb of the northernmost synformal structure is contained within the composite profile of fig. 3. folding of the aqd was observed by henderson (1969) and later elaborated upon by passchier et al. (1997) and van gool et al. (1999). at the localities outside the shear zone where the aqd can be seen in contact with the structurally underlying siliceous gneisses (localities 01702, arfersiorfik inner arfersiorfik b a orlerf ik sarfaarsuk nuerso rfik 12702 01703 03706 13705 13704 01704 15705 15703 15702 51° 68° 50°30' 51°30' major outwash plains fault (inferred) structural trend line 10 km polonia gletscher highly deformed, chemically altered and pegmatite-invaded part of the aqd ('pencil gneiss'), mapped and inferred arfersiorfik quartz diorite (aqd), mapped and inferred supracrustal schist units, mapped and inferred siliceous gneisses (probably orthogneisses), mapped and inferred nuussuaqtunertoq inland ice 18702 05702 aku nn aa q 05701ks fig. 2. simplified geological map of the inner arfersiorfik region, showing the authors’ 2002 field work along full-coloured coasts (see also fig. 1). compare the mapped geology with the magnetic patterns on fig. 7, shown at the same scale. red lines, structural trend lines mainly from aeromagnetic data. ks, kangimut sammisoq (near map centre). five digit numbers, locality numbers mentioned in the text. two antiformal folds east of the inferred fault along inner arfersiorfik are outlined by the aqd contact overlying siliceous gneisses north of localities 05702 and 03706. the axial trace of the intervening synformal fold meets the coast near loc. 01703. tent symbol, position of camp during 2002 field work. inland geology near nuersorfik from van gool & marker (2004). 127020170201703 (cumulate) 02701 02702 08703 08704 18702 metadolerite locality aqd ba sse nnw (31 km south) 1 km 'hornblende gneiss rich in microcline' of henderson (1969): lineated and fine grained with abundant pegmatite sheets nordre strømfjord shear zone granitic orthogneiss amphibolites mafic enclaves hornblendebearing orthogneiss gneisses with abundant inclusions massive coarse-grained aqd homogeneous gneisses with few inclusions neosome increases three metadolerites 4th supracrustal unit 5th supracrustal unit 3rd supracrustal unit 2nd supracrustal unit 1st supracrustal unit fig. 3. schematic profile a–b through the nordre strømfjord shear zone in the inner arfersiorfik from akunnaaq in the south to sarfaarsuk in the north (figs 2, 7). localities and lithologies have been projected onto the profile plane along the strike direction of the shear zone. locality numbers mentioned in the text appear beneath the profile. for colour code, consult caption to fig. 2. 149 01704, 03706, 05701, figs 2, 3), the contacts are all tectonised, and with a well-developed, folded planar fabric. original intrusive features have been obliterated except at loc. 05702 (fig. 2), where xenoliths of pelitic schist can be seen within the aqd, clearly indicating intrusion of the quartz diorite into a pre-existing metamorphosed sedimentary sequence. from loc. 01702 (fig. 3) and towards the north, siliceous gneisses become progressively more deformed and contain an increasing amount of granitic material. for a distance of almost 1 km to the south of the first supracrustal schist unit, the gneisses display a subvertical planar fabric and contain abundant layers of leucocratic, in places pegmatitic, material. these gneisses appear highly strained. the transition within these gneisses, from 01701 to 02702 (fig. 3), is the best exposed transition zone developed in siliceous gneisses in the area that we have visited. the structural development of these gneisses is illustrated in fig. 4a–c. the first supracrustal unit (fig. 3) comprises a predominant lithology of mafic schists with calc-silicate and ultramafic lenses. pelitic schists are subordinate. at locality 08703, a second thin supracrustal unit of pelitic schist (thickness 25 m) occurs along the boundary to the aqd. the gneisses to the north of the first supracrustal unit contain abundant layers of homogeneous amphibolite, a c b b d fig. 4. siliceous gneisses between localitiess 17701 and 02702 (figs 2–3) illustrating progressive deformation towards the first supracrustal unit. a: coarsely banded biotite gneiss, loc. 17701 (knife 10 cm long). b: finely banded gneiss with generally steeply oriented foliation, but apparently variable deformation intensity at outcrop scale, loc. 02701 (compass 10 cm long). c: extreme planar fabric and fine-scale layering in siliceous gneiss near the first supracrustal schist unit, loc. 02702. d: detail from the same locality with lenticular ultramafic inclusions (knife 10 cm long). 150 with locally preserved discordant relationships to the host gneiss structures. these discordant amphibolites were undoubtedly originally dolerite dykes. farther to the north, the gneisses are nearly devoid of amphibolites and are characterised by wavy and folded migmatitic structures without the pronounced planar fabric seen in the gneisses adjacent to the first supracrustal unit. the contact between the second supracrustal unit and the aqd is vertical. coarse-grained segregations of granitic pegmatite in the schist appear to be smeared out and pulled apart. the aqd near this contact is syntectonically recrystallised and displays layering on a centimetre scale. the aqd farther to the north of the contact becomes relatively coarse grained and in places preserves undeformed magmatic structures (see a later section with more details on the petrology of the aqd). these relatively coarse-grained aqd rocks become progressively tectonised towards the north, developing a very pronounced ltectonite fabric. pegmatite sheets also become abundant, and the aqd itself becomes progressively more leucocratic. henderson (1969 p. 7) describes this transition with these words: “in the inner part of arfersiorfik a zone of basic gneiss of a type not found elsewhere in the area was encountered. this is a hornblende gneiss rich in microcline. it contains numerous pegmatites”, and further, “a large folded quartz diorite body dominates the inner part of arfersiorfik ... on its northern side it is in contact with the hornblende gneiss rich in microcline mentioned previously” (henderson 1969, p. 11). we interpret the “hornblende gneiss rich in microcline” as the northernmost, most highly deformed (and most chemically altered) equivalent of the much less deformed and more igneous appearing aqd to the south of the shear zone. nevertheless, the aqd to the south of the shear zone is in most places also a tectonite; a mineral lineation is nearly always present in the rock. textures viewed on sections perpena b fig. 5. arfersiorfik quartz diorite. a: coarsegrained metamorphosed quartz diorite with pervasive linear fabric. a minor dextral shear zone deforms the linear fabric (compass 10 cm long). b: extreme linear fabric developed in aqd close to its northern boundary (‘pencil gneiss’; knife 20 cm long). 151 dicular to the lineation appear igneous, while a pronounced tectonic lineation is seen in other orientations (fig. 5). the transition from the more magmatic appearing aqd to the “hornblende gneiss rich in microcline” is gradual, and the continuity of these rocks is evident on either side of arfersiorfik. between the third and fourth supracrustal units siliceous gneisses of varying heterogeneity and apparent degree of shear zone deformation occur, all with a vertical planar structure. basic inclusions are common in this gneiss unit. the fourth supracrustal unit as exposed on the eastern shore of orlerfik is a lithologically variable assemblage of pelitic schists and gneisses, amphibolitic schists, and calcsilicate gneiss. overall, the rocks appear highly deformed. the pelitic rocks typically contain numerous layers and lenses of pegmatitic material up to 8 m thick. cordierite was observed in the schists, and a lens of tourmalinite was also found. near the inland ice, at loc. 15703 (fig. 2), a centimetre-thick marble layer, the only occurrence of marble seen during field work, was observed in a pelitic schist thought to be the merged equivalent of the third and fourth supracrustal units at orlerfik. the gneisses between the fourth and fifth supracrustal a b fig. 6. pelitic and mafic schists from supracrustal unit near the inland ice visited during helicopter reconnaissance, loc. 15705 (fig. 2). a: two-mica sillimanite schist with abundant ‘winged’ feldspar blasts indicating sinistral sense of shear. long dimension of photo c. 75 cm. b: amphibolitic schist with boudins of garnet-bearing ultramafic rock (hammer c. 50 cm long). 152 units are quartzofeldspathic. a rather homogeneous, granitic orthogneiss 1 km wide occurs immediately to the north of the fourth supracrustal unit. it contains a few mafic inclusions which appear to be metadolerites. this gneiss, although foliated, has not developed the typical shear zone fine layering, and it is clearly an orthogneiss. towards the north, quartzofeldspathic gneisses with numerous granitic layers and bands form the boundary towards the fifth supracrustal unit. the finely banded shear zone planar structure is developed in places. the gneisses between the fourth and fifth supracrustal unit thus appear to be progressively deformed towards the north. the fifth supracrustal unit is well exposed at orlerfik and along the western shore of the nuussuaq peninsula. on nuussuaq, this unit forms an elongate ridge of strongly rusty weathering rocks. in fresh exposures in the tidal zone, the rusty weathering rock can be seen to be a fine-grained graphiteand sulphide-bearing darkish grey gneiss. in this unit, finely banded amphibolitic schists are prominent, along with pelitic sillimanite-bearing schists. calc-silicate rocks are seen as well, as also described by henderson (1969). pegmatitic layers are abundant. we were unable to confirm the marble layer occurring on the western half of the nuussuaq peninsula on henderson’s (1969) map. we found, instead, a complex of light-coloured, in some places greenish, anthophyllite-phlogopite-diopside-bearing metamorphosed ultramafic rocks. it appears that all the supracrustal units, although structurally continuous, change lithologically along strike, consisting of various combinations of pelitic schists and gneisses, graphite-sulphide gneisses, calc-silicate rocks, mafic schists and ultramafic lenses. in all the supracrustal units deformation appears to be intense. both pelitic and mafic rocks are finely schistose (fig. 6). ultramafic and skarn lenses may be rather massive and homogeneous, but they are small in size and volumetrically minor. overall, the supracrustal units are schistose throughout, and differ in this way from the intervening siliceous gneisses, which are heterogeneously deformed on a 100 m scale. to the north of the fifth supracrustal unit only scattered exposures of gneiss are found at orlerfik. along several localities near the northern boundary of the shear zone, at orlerfik and along the south coast of tunertoq, a homogeneous, hornblende-bearing orthogneiss is encountered. in places it is reminiscent of the aqd. farther north of the shear zone in the area around sarfaarsuk, gneisses with a complex history occur, as indicated by several generations of granitic and pegmatitic veins and layers, and mafic and ultramafic layers and lenses. at locality 12702 (fig. 2), a number of mafic layers are clearly folded metadolerites, and at this locality it can be seen how new layering develops concomitantly with deformation. along this layering the thin metadolerites can be seen to be displaced, and it seems evident that new layering develops in the siliceous gneisses at small angles to the older layering. without the metadolerites it would be impossible to discriminate between new and old layering. aeromagnetic data the entire nordre strømfjord shear zone and adjacent areas are covered by regional aeromagnetic data (figs 1b, 7) acquired during the project aeromag 1999 (see text to fig. 1b and rasmussen & van gool 2000). the aeromagnetic map proved very useful for delineation of lithologies and structures in the field, and in general there is a very clear correlation between observed surface geology and the aeromagnetic anomaly patterns. the arfersiorfik quartz diorite is associated with a variable magnetic signature. the best exposed part of the aqd is located in inner arfersiorfik just north of polonia gletscher (fig. 7). this part is denoted the core zone of the aqd complex and is characterised by a distinct positive anomaly (300–490 nt). here an unusually large amount of magnetite was found, both as millimetre-scale magnetite needles and as magnetite grains 1–3 cm large. a pronounced decrease in the magnetic intensity of the aqd is observed just north-west of the core zone. the magnetic signature of the aqd changes towards the shear zone, where the fine-grained, pegmatite-bearing and strongly linear aqd (the ‘pencil gneiss’) that forms the northernmost 1–2 km of the body exhibits strong negative anomalies. these changes in magnetic signature appear to reflect metamorphic and metasomatic processes associated with the shear zone development. the demagnetising effects reflect the formation of titanite and biotite at the expense of fe-ti-oxides, which is synchronous with shear zone evolution. detailed thin section observations show that the modal abundances of titanite and magnetite/ilmenite are strongly negatively correlated. comparison of titanite-rich samples and magnetite/ilmenite-rich samples with the vertical gradient of the magnetic anomaly field (obtained from the map shown in fig. 7) shows that the titanite-rich samples consistently occur within negative changes in the gradient, while the magnetite/ilmenite-rich samples occur within positive gradients. further work is underway to evaluate the reliability of this correlation. in general, the supracrustal schist units are associated with negative anomalies in the total magnetic field intensity. within the shear zone, the schist units are characterised by pronounced, linear short-wavelength negative ano153 malies (–150 to –300 nt). at kangimut sammisoq (ks on fig. 7) and towards the west-south-west the aqd and juxtaposed schists form a coherent negative anomaly, which can be traced for a considerable distance towards west-south-west into the agto map sheet area (olesen 1984). within the latter area the anomaly reflects supracrustal rocks only. the negative anomalies of the supracrustal schists probably reflect the absence of fe-ti oxides and the abundance of iron sulphides and graphite, which yields a weak magnetic response. in general, gneiss lithologies in the area are expressed as positive magnetic anomalies. because of their greater volume, the gneisses are associated with broader wavelength anomalies than the supracrustal units. however, the shape of the anomalies associated with the gneiss units within the shear zone are of short wavelength and elongate (fig. 7). where a significant strain gradient in the siliceous gneisses is observed, e.g. northwards from akunnaaq along the western shore of inner arfersiorfik (from loc. 01702 to 02702, fig. 3), this transition is not visible in the magnetic field (a on fig. 7). in general, it appears that even large changes in strain adjacent to and within the shear zone do not have a clear magnetic signature in the siliceous gneisses. while the magnetic lows associated with supracrustal units are in general persistent along strike, the highs associated with the siliceous gneisses in places outline lenticular structures, the most obvious example being the magnetic high which can be seen to coincide with the orthogneisses between the fourth and fifth unit (b on fig. 7). the general map structure although we have not carried out regional mapping, some map scale observations follow from our work. from the inland ice and for a distance of at least 35 km towards the south-west (i.e. in the area between the inland ice and the bend of arfersiorfik towards south-east west of nuussuaq), the aqd is bounded to the north by the shear zone, and probably with a continuous screen of supracrustal schists along its northern boundary. towards this bounding schist, the aqd becomes progressively deformed, developing a pronounced linear fabric and undergoing a number of [nt] 524 207 144 104725032186–6 –15 –24 –32 –40 –48 –56 –64 –71 –79 –86 –93 –100 –108 –115 –122 –129 –137 –144 –151 –159 –167 –175 –184 –192 –202 –214 –227 –245 –275 –402 ks e b a c 51°30' 50°30' 68° d 51° inland ice polonia gletscher inner arfersiorfik aku nn aa q orlerf ik arfersiorfik nuerso rfik qaamasuuoq sarfarsuuk nuussuaq tunertoq 10 km fig. 2 fig. 7. total magnetic field intensity map of the study area. a, akunnaaq; ks, kangimut sammisoq; b, magnetic high coinciding with the orthogneisses between the fourth and fifth supracrustal units; c–d, lithological boundary between siliceous gneisses and supracrustal schists; e, possible magnetic expression of merged fourth and fifth supracrustal units. the position of the camp site used during the field work is shown with a tent symbol. compare the magnetic patterns with the mapped geology on fig. 2 (white frame), shown at the same scale. nt: nanotesla, a unit of magnetic flux density. 154 mineralogical changes, described below. these mineralogical changes are reflected in a change in magnetic signature. the change is less dramatic to the east of arfersiorfik than it is to the west of the fjord, and this signature can be followed from the fjord to the inland ice and beneath the ice for at least another 30 km. based on aeromagnetic data covering the entire greenland including the inland ice, nielsen (2004) and nielsen & rasmussen (2004) suggest a 125 km eastward extension of the aqd beneath the ice. the outcrop pattern and orientations of the boundary between the aqd and the structurally underlying schists and gneisses are in accord with a structural model involving relatively open folding, as marked on fig. 2. on henderson’s (1969) map, a ‘tongue’ of siliceous gneisses projects from the eastern shore of the fjord 20 km eastwards into the aqd. during helicopter reconnaissance by us, this eastward-reaching gneiss was nowhere to be seen. along the shores of arfersiorfik, there is an antiformal closure of aqd over siliceous gneisses with an intervening schist screen, but this cannot extend inland for more than 2 km. the entire triangular area bounded by arfersiorfik, the shear zone, and the inland ice is most likely only underlain by aqd. the map pattern west of inner arfersiorfik differs from the pattern described above. along the shoreline south of kangimut sammisoq the southern boundary of the aqd is vertical, and south of the boundary, gneisses and the schists of the first supracrustal unit are also vertically oriented and possess a well-developed planar shear zone fabric. we interpret this difference in map pattern as reflecting intersections with the shear zone boundary at two different levels. petrology arfersiorfik quartz diorite the arfersiorfik quartz diorite was first mapped in the study area by henderson (1969), who noted that it varies from a homogeneous, coarse-grained igneous rock in its central region which lacks evidence of deformation, to a compositionally banded gneiss near its margins. further work by kalsbeek et al. (1984, 1987) and kalsbeek (2001) supported the general descriptions of henderson and provided a considerable body of evidence regarding the age and chemical composition of the body. our observations expand on these earlier results. within the ‘greater aqd igneous complex’ we have observed a number of lithologies indicative of a complex igneous history (fig. 8). at loc. 01703 (fig. 2) we observed a locally developed layered cumulate sequence (fig. 8a) in gradational igneous contact with quartz diorite that contains blocks of more mafic igneous rocks. the cumulate sequence consists of numerous cycles of pyroxene-olivine cumulate layers grading into plagioclase-pyroxene anorthositic layers. this rhythmic layering grades upward into massive dioritic and quartz-dioritic rocks that contain plagioclase-porphyritic mafic xenoliths. at the same structural level, but farther north at loc. 08704, mafic enclaves (fig. 8b) occur within leucocratic quartz diorite. in places, these enclaves contain plagioclase-porphyritic xenoliths similar to those observed at loc. 01703. we interpret these enclaves to be evidence of mixing of magmas, similar to that reported for other igneous bodies (e.g. gagnevin et al. 2004; healy et al. 2004; janousek et al. 2004). the presence of both cumulate rocks and mixed mafic magmas at approximately the same structural level suggests that these sites represent the lower levels of the igneous complex that makes up the aqd suite. at loc. 13705 clear evidence of at least two separate magmas in an undeformed state is preserved, with a later diorite exhibiting well-developed chilled margins against an earlier quartz diorite (fig. 8d). whether this implies that the aqd is actually a collection of numerous smaller intrusions or a major igneous body with local, small and rare later intrusive phases can only be resolved with more detailed field work. at loc. 18702 (fig. 2) we have also observed a finely banded rock, apparently developed by extreme deformation of a sequence of pillow lavas (fig. 8d). this latter occurrence was found close to the northern boundary of the aqd and may represent the upper part of the ‘greater aqd igneous complex’. within the thin sheet of aqd to the south of the localities that we examined, an occurrence of metavolcanic rocks was reported by manatschal et al. (1998), which supports the interpretation that exposures of the aqd suite represent a variety of deep to near-surface original positions. in thin section, the mineralogy of the aqd suite is invariably modified by metamorphic recrystallisation. textural features suggest that the aqd originally consisted of hornblende and pyroxene diorites and quartz diorites. in those rocks with primary pyroxenes, the orthoand clinopyroxenes are partially to completely recrystallised to intergrowths of plagioclase, hornblende, quartz, magnetite, ilmenite and sometimes biotite. in instances where the pyroxene is completely replaced by amphibole, the amphibole is densely filled with inclusions of magnetite/ ilmenite. in the hornblende diorites and quartz diorites the original igneous amphiboles are generally inclusionfree and tend to have more brownish pleochroic colours. in some cases, plagioclase crystals exhibit well-preserved 155 igneous oscillatory zoning, but usually this is partially to completely replaced by metamorphic zoning. cummingtonite-hornblende intergrowths are relatively common and are identical to similar features described from dioritic rocks affected by amphibolite facies metamorphism in other parts of the world, e.g., tanzania (haslam & walker 1971), usa (new hampshire, brady 1974; the grand canyon, clark 1978), and australia (new south wales, stephensen & hensel 1979). recrystallisation of the aqd occurred within middle to upper amphibolite facies, as indicated by the hornblende-biotite-plagioclase ± garnet association, the instability of orthoand clinopyroxene with amphibole, and the development of cummingtonitehornblende intergrowths. the chemical variability of the aqd suite (fig. 9) is typical of that of other calc-alkaline igneous complexes, as noted previously by kalsbeek et al. (1987) and kalsbeek (2001) who demonstrated similarities between these rocks and the sierra nevada batholith. including their data, we have expanded the comparison to that of the cascades of washington state, usa, the mt lassen complex of california, and the southwest pacific island arc a b b c d fig. 8. extent of lithological variation within the ‘greater arfersiorfik quartz diorite complex’. a: igneous layering at loc. 01703 (fig. 2), where an assemblage of undeformed cumulates and related rocks from the aqd suite is set in a host of tectonised (lineated) aqd (pencil c. 15 cm long). b: igneous textured aqd with hornblende-rich pods indicating fractionation or magma mingling. observe light grey xenolith within hornblende rich phase immediately north-east of centre of photo (fist for scale). c: dark coloured igneous rock (left) chilling against lighter coloured rock (right), loc. 13705 (fig. 2; compass 10 cm long). d: part of a pegmatite sheet (right) within a banded amphibolitic rock (left) interpreted as strongly deformed pillow lava (hammer handle c. 50 cm long). 156 system. the compositional range of the aqd suite is large, varying from ultramafic (within the cumulate complex) to quartz-dioritic and quartz-monzonitic. the compositional evolution of the magma appears to have been controlled by crystal fractionation of olivine and/or pyroxenes when the cumulates were forming, but eventually reaching the stage where hornblende fractionation dominated the chemical evolution of the magma (fig. 10). the one sample that deviates from these trends is labelled m-aqd. this is a fine-grained biotite-plagioclase-quartz mylonite produced by local granulation of coarse-grained aqd. further study of this sample is underway to understand the controls on this alteration process. we noted that in the near vicinity of the nordre strømfjord shear zone chemical modification of the aqd appears to be significant. development of k-feldspar-rich mineral assemblages is common, and often accompanied by the association calcite-titanite. these rocks occur in the same region in which henderson (1969) noted the presence of k-feldspar-rich gneisses. the magnitude, extent and significance of these chemically unusual rocks are under investigation. metasomatism at the aqd margin and minor shear zones within the aqd were described by kalsbeek et al. (1987). included in figs 9, 10 are four analyses (labelled lb) from samples of large boulders in outwash from glacial deposits, collected several hundred metres east of our camp (see fig. 2). in the field, the samples bear striking similarity to aqd lithologies. as is evident from the figures, these rocks are chemically indistinguishable from the aqd suite, supporting evidence from aeromagnetic data discussed above that the aqd extends under the ice for some distance. late garnet granite the southernmost samples collected (loc. 13704, fig. 2) are from a site at which peraluminous garnet granite exhibits classic intrusive relationships into granodioritic to quartz-dioritic gneisses of the aqd. this granite contains stoped blocks of gneisses within the upper 30 m of its upper boundary zone (fig. 11), and dykes and sills from the granite invade the gneiss. the granite itself contains lb lb mt lassen new zealand and tonga-kermadec aqd suite aqd (kalsbeek 2001) mt stuart batholith fe2o3 + feo mgona2o + k2o m-aqdm-aqd skaergaard fig. 9. afm diagram comparing arfersiorfik quartz diorite chemistry (including the large boulder samples labelled ‘lb’) to that of other calc-alkaline systems. the trend for the skaergaard igneous complex is shown for comparison. mt stuart data from erikson (1977). all other trends from best (1982). the sample labelled m-aqd is the mylonitised quartz diorite sample mentioned in the text. lblb aqd cumulate suite aqd plutonic suite aqd (kalsbeek 2001) al-na-k fe + mgca m-aqdm-aqd cordierite biotite anorthite clinopyroxene hornblende fig. 10. acf atomic triangular diagram showing the chemical variation in the arfersiorfik quartz diorite based on our data (including the boulder samples lb and data reported by kalsbeek 2001). red symbols, samples collected within cumulate localities. blue and green symbols, collected by us and kalsbeek (respectively) from the non-layered igneous complex. note the kink in the trend of chemical evolution and the overlap between the non-layered and cumulate rocks of the arfersiorfik quartz diorite. red arrowed line, the trend of magmatic evolution implied by the cumulate samples, reflecting fractionation of some combination of olivine and/or orthopyroxene with clinopyroxene in the most primitive cumulate suite. blue arrowed line, that suite of samples for which chemical variability can be ascribed to hornblende fractionation. yellow box: outline of the range of likely hornblende compositions. m-aqd, mylonitised quartz diorite sample mentioned in the text. 157 delicate, primary muscovite rosettes and myrmekitic intergrowths with no evidence of deformation. the intrusive relationships and the preservation of delicate primary igneous structures suggest that the granite invaded postkinematically, and that it could provide excellent samples for establishing a minimum age for the time of cessation of tectonic activity in the region. late stage, low-grade metamorphism a suite of samples from the aqd in the region south of the nordre strømfjord shear zone contain postkinematic mineral assemblages of the prehnite-pumpellyite and lower greenschist facies overprinting the higher grade metamorphic mineral assemblages that define the amphibolite facies metamorphism. this late-stage, low-grade metamorphism is associated with brittle fracturing and veining, and local piemontite and base metal mineralisation. the timing of this low-grade metamorphism remains unresolved. strain variation within the nordre strømfjord shear zone earlier work within the agto map sheet (fig. 1; bak et al. 1975b; sørensen 1983) concluded that the nordre strømfjord shear zone reflects sinistral transcurrent movement. it was also concluded from the fanning of planar structures that the shear zone is wedge-shaped within the agto map sheet. the present work has shown that in the inner part of arfersiorfik such fanning is not obvious, perhaps because the zone here is only half as wide as near the coast in the agto map sheet area. the planar structures within the shear zone are vertical (fig. 3) with an extremely limited variation of dips (fig. 12), and linear structures are horizontal. the justification for drawing the schematic profile (fig. 3) with vertical lithological boundaries throughout the shear zone is evident from the stereograms of fig. 12. the horizontal orientation of linear structures is also in accordance with a horizontal movement direction and with the orientation of linear structures in the western part of the shear zone (bak et al. 1975b). the sinistral shear sense, already inferred by bak et al. (1975b) and further discussed and analysed by sørensen (1983), is obvious from the large-scale anticlockwise rotation seen in fig. 11. peraluminous garnetand muscovite-bearing granite at loc. 13704 (fig. 2) intruding into arfersiorfik quartz diorite and interpreted as the roof zone of a pluton with rafts of country rock and sheets of granite above. looking north-north-east. height of section c. 150 m. 158 the aeromagnetic data (figs 1, 7). small-scale shear sense indicators confirm this pattern (fig. 6). they do not occur abundantly, and as the large-scale features of the entire shear zone from the coast to the inland ice leave no doubt about the sinistral nature of movement, we have made no systematic investigations of small-scale shear sense indicators. it was suggested by sørensen (1983, fig. 14) that the shear zone between arfersiorfik and the inland ice could be divided into an central part bounded by two marginal zones. he furthermore speculated (p. 3432) that the strain profile at the inland ice was suggestive of superposition of two mechanisms, one responsible for most of the deformation in the marginal zones, and another responsible for the major part of the strain in the central part of the shear zone. the rotation on which this suggestion was based is also evident, for example, on the aeromagnetic data between labels c and d on fig. 7, and from there b a c d e f n = 85 n = 25 n = 52 n = 12 n = 37 n = 23 c.i. = 3.0 % / 1 % area south of the shear zone within the shear zone north of the shear zone planar structures linear structures fig. 12. stereograms of planar and linear structures from within and outside the nordre strømfjord shear zone. while linear structures (d–f) outside and within the shear zone do not differ in orientation, planar structures (a–c) within the shear zone are seen to be extremely concentrated as opposed to the wider spread of planar structures along girdles perpendicular to the linear structures outside the zone. equal area lower hemisphere projections. contour interval = 3% of plotted points per 1% of stereogram area. 159 farther towards the west-south-west. our field work demonstrates that this line is in fact a lithological boundary between siliceous gneisses and supracrustal schists. a corresponding boundary may be placed at the boundary between the ‘pencil type’ aqd and the schist that we have termed the third supracrustal unit on fig. 3, and thus like the northern line of rotation coincides with a lithological boundary. throughout the shear zone all the supracrustal units are highly and penetratively deformed, whereas the intervening orthogneisses are heterogeneously deformed. furthermore, where visited by the authors, the boundaries to the central part of the shear zone do in fact correspond to lithological boundaries involving supracrustal schists. in addition, the intervening siliceous gneisses, as judged from the aeromagnetic map, may in places form augenlike features, as for example, immediately north of e on fig. 7. the eastern tip of this positive magnetic anomaly was visited by helicopter (loc. 15702). immediately east of this locality the fourth and fifth supracrustal units can be seen to be separated by a few tens of metres of siliceous gneiss, a unit which is almost 2 km wide along orlerfik 15 km to the west. to the east of this locality, a wide negative magnetic anomaly can be seen to underlie unexposed ground (e on fig. 7). we interpret this as the magnetic expression of the merged supracrustal units four and five. our work in arfersiorfik therefore suggests a model for strain variation that is fundamentally different from the model developed by sørensen (1983) for the western, lower crustal and most high-grade part of the shear zone. where the latter model suggests continuity of lithologies from wall to wall of the shear zone, our observations in arfersiorfik indicate that adiscontinuitiesy in strain may be associated with boundaries between schists and orthogneisses. it is furthermore suggested by the aeromagnetic data that the intervening siliceous gneisses may form augen-like features. the change in deformation mechanism suggested by sørensen (1983) may therefore reflect ductility contrasts between the supracrustal schist units and the siliceous gneisses and not the superposition of two different deformation mechanisms. mineral occurrences two dominating types of mineral occurrences have been found in the inner part of arfersiorfik: (1) occurrences of disseminated to massive pyrrhotite in supracrustal mica schist units and (2) occurrences of iron sulphides in the contact/fault zones between the aqd and juxtaposed supracrustal units. the former type is associated with strongly foliated amphibolite, biotite ± graphite ± sillimanite schist and paragneiss. the pyrrhotite is found either as horizons up to 0.5 m thick or as lenticular bodies or pods up to 1 × 2 m in outcrop size. this type of occurrence is only found in the supracrustal rock units within the shear zone. similar occurrences exist west of the investigated area all the way to the outer coast just north of the mouth of nordre strømfjord. this occurrence type is interpreted as syn-genetic volcanic-associated exhalative sulphide deposited in a volcanic arc environment. the highest analytical values of gold and base metals obtained from samples of disseminated sulphides in pelitic schist are the following figures: 38 ppb au, 45 ppm as, 497 ppm zn, 3813 ppm cu and 456 ppm ni. all samples have elevated mn concentrations, up to 2.9 wt% mn. the second type of mineral occurrence is widespread in the inner part of arfersiorfik, both within and outside the shear zone. the zones hosting the sulphides are often associated with alteration, and in many cases a high degree of silicification is observed. it is unknown if this type is strictly epigenetic, caused by hydrothermal activity along the contact or in the fault zones between the aqd and the minor supracrustal units, or if it is a product of hydrothermal remobilisation of syngenetic iron sulphides in the supracrustal rock units. for more details and description of mineral occurrences in central west greenland see steenfelt et al. (2004) and stendal et al. (2004). conclusions our investigations of the nordre strømfjord shear zone in the eastern part of arfersiorfik suggest that deformation and displacement at this crustal level are primarily a function of lithology, with large displacements effected by ductile flow in the supracrustal sequences. the large mass of arfersiorfik quartz diorite between arfersiorfik and the inland ice is bordered to the north by the shear zone, probably with a supracrustal schist unit along this boundary. the latter is indicative of the extreme deformation that can be taken up in the supracrustal rocks. towards this boundary the tectonite fabric of the aqd becomes further accentuated, and near its northern boundary the aqd differs both chemically, mineralogically and structurally from the aqd to the south of the shear zone. we have both found cumulate rocks originating from near the base of the original magmatic aqd complex and metavolcanic rocks interpreted as belonging to its top part, whose less deformed equivalents are most probably covered by the inland ice. the chemical variation of the aqd suite is wide and comparable to that of cenozoic calcalkaline igneous suites. 160 acknowledgements concise reviews by graham leslie and chris pulvertaft and financial support by the carlsberg foundation are gratefully acknowledged. the air greenland helicopter pilot and the crew on board m/s milandt (aasiaat) are thanked for efficient transport and practical help. chemical analyses were carried out by xrf on glass discs at the department of earth sciences, university of aarhus, denmark. references bak, j., grocott, j., korstgård, j.a., nash, d., sørensen k. & watterson, j. 1975a: tectonic implications of precambrian shear belts in western greenland. nature 254, 566–569. bak, j., korstgård, j.[a.] & sørensen, k. 1975b: a major shear zone within the nagssugtoqidian of west greenland. tectonophysics 27, 191–209. best, m.g. 1982: igneous and metamorphic petrology, 630 pp. san francisco: w.h. freeman. brady, j.b. 1974: coexisting actinolite and hornblende from west-central new hampshire. american mineralogist 59, 529–535. clark, m.d. 1978: amphibolitic rocks from the precambrian of grand canyon: mineral chemistry and phase petrology. mineralogical magazine 42, 199–207. connelly, j.n., van gool, j.a.m. & mengel, f.c. 2000: temporal evolution of a deeply eroded orogen: the nagssugtoqidian orogen, west greenland. canadian journal of earth sciences 37, 1121–1142. erikson, e.h. jr. 1977: petrology and petrogenesis of the mount stuart batholith – plutonic equivalent of the high-alumina basalt association?. contributions to mineralogy and petrology 60, 183–207. escher, a. 1971: geological map of greenland, 1:500 000, søndre strømfjord – nûgssuaq, sheet 3. copenhagen: geological survey of greenland. gagnevin, d., daly, j.s. & poli, g. 2004: petrographic, geochemical and isotopic constraints on magma dynamics and mixing in the miocene monte capanne monzogranite (elba island, italy). lithos 78, 157–195. hanmer, s., mengel, f., connelly, j. & van gool, j.[a.m.] 1997a: significance of crustal-scale shear zones and synkinematic dykes in the nagssugtoqidian orogen, sw greenland: a re-examination. journal of structural geology 19, 59–75. hanmer, s., mengel, f., connelly, j. & van gool, j.[a.m.] 1997b: significance of crustal-scale shear zones and synkinematic dykes in the nagssugtoqidian orogen, sw greenland: a re-examination. 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(ed.): fluid movements – element transport and the composition of the deep crust, 171–185. amsterdam: kluwer academic publishers. steenfelt, a., stendal, h., nielsen, b.m. & rasmussen, t.m. 2004: gold in central west greenland – known and prospective occurrences. geological survey of denmark and greenland bulletin 4, 65–68. stendal, h., nielsen, b.m., secher, k. & steenfelt, a. 2004: mineral resources of the precambrian shield of central west greenland (66° to 70°15′n). part 2. mineral occurrences. danmarks og grønlands geologiske undersøgelse rapport 2004/20, 212 pp. stendal, h., secher, k. & frei, r. 2006: 207pb-206pb dating of magnetite, monazite and allanite in the central and northern nagssugtoqidian orogen, west greenland. in: garde, a.a. & kalsbeek, f. (eds): precambrian crustal evolution and cretaceous–palaeogene faulting in west greenland. geological survey of denmark and greenland bulletin 11, 101–114 (this volume). stephensen, n.c.n. & hensel, h.d. 1979: intergrown calcic and femg amphiboles from the wongwibinda metamorphic complex, n.s.w., australia. canadian mineralogist 17, 11–23. van gool, j.a.m., kriegsman, l.m., marker, m. & nichols, g.t. 1999: thrust stacking in the inner nordre strømfjord area, west greenland: significance for the tectonic evolution of the palaeoproterozoic nagssugtoqidian orogen. precambrian research 93, 71–86. van gool, j.a.m., connelly, j.n., marker, m. & mengel, f.c. 2002: the nagssugtoqidian orogen of west greenland: tectonic evolution and regional correlations from a west greenland perspective. canadian journal of earth sciences 39, 665–686. van gool, j.a.m. & marker, m. 2004: geological map of greenland, 1:100 000, ussuit, 67 v.2 nord. copenhagen: geological survey of denmark and greenland. __________________________________________________________________________________________________________________________________________________________________________________ manuscript received 5 november 2004; revision accepted 1 november 2005 162 research article sandersen et al. 2023: geus bulletin 53. 8351. https://doi.org/10.34194/geusb.v53.8351 1 of 14 transport of nitrate-containing groundwater to coastal areas through buried tunnel valleys, denmark peter b.e. sandersen*1 , hyojin kim2 , rasmus jacobsen2 , jesper b. pedersen3 , birgitte hansen2 1department of near-surface land and marine geology, geological survey of denmark and greenland, aarhus, denmark. 2department of geochemistry, geological survey of denmark and greenland, copenhagen, denmark. 3hydrogeophysics group, department of geoscience, aarhus university, aarhus, denmark abstract nitrogen impact on the aquatic environment, including coastal areas, is too high in many countries worldwide, particularly in areas with intensive agriculture. efficient mitigation initiatives demand that important pathways and the fate of nitrate in the hydrological cycle are known. in this study, we focus on groundwater nitrate contamination in two near-shore catchment areas in north-west denmark. groundwater in the area is mainly located in buried tunnel valleys, which are subsurface structures eroded by meltwater during pleistocene glaciations in former glaciated areas. groundwater samples from the aquifers inside the buried valleys reveal the presence of up to 120 mg/l nitrate down to 10 m below sea level and about 1 km down from the stream outlet towards the coast. we interpret the complex tunnel-valley infill to be responsible for the spatial heterogeneity of the groundwater geochemistry, where sandy geological windows create localised hydraulic pathways and complex redox structures. groundwater and stream water chemistry in the study area clearly demonstrate the role of groundwater in nitrate transport within the catchment as well as the direct pathway to the coast bypassing the stream and riverine systems. our results show that the buried tunnel valleys potentially contribute to submarine groundwater discharge and therefore could be responsible for a hitherto unaccounted input of nitrogen to the marine environment. *correspondence: psa@geus.dk received: 17 apr 2023 revised: 31 july 2023 accepted: 10 aug 2023 published: 27 nov 2023 keywords: submarine groundwater discharge, redox modelling, buried tunnel valleys, nitrate in groundwater, nitrogen flux to the marine environment abbreviations a.s.l.: above sea level b.l.s.: below land surface b.s.l.: below sea level c–q: nitrate concentration–discharge doi: depth of investigation geus: geological survey of denmark and greenland n: nitrogen ofg: offshore fresh groundwater sgd: submarine groundwater discharge tem: transient electromagnetic ttem data: towed transient electromagnetic geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: julian koch (geus, denmark) reviewed by: theis r. andersen (via university college, denmark) and erwin racasa (university of rostock, germany) funding: see page 12 competing interests: see page 13 additional files: none 1. introduction eutrophication of sea waters is a global phenomenon, due to nitrogen (n) enrichment that mainly stems from inland riverine systems (e.g. defra 2016). anoxic dead zones, which are considered a key stressor on marine ecosystems, can be the consequence (diaz & rosenberg 2008). nitrogen pollution to coastal zones around the world originates from different terrestrial sources, such as waste water, atmospheric fixation and deposition, livestock manure and fertiliser application (erisman et al. 2011). the implementation of danish national action plans since the mid-1980s for the reduction of n losses has successfully reduced the impact of n on the aquatic environment (dalgaard et al. 2014; hansen et al. 2017). however, denmark still has a relatively high level of agricultural n loss, surpassed only by a few areas in europe, such as in the netherlands and the po valley in italy (de vries et al. 2011). to achieve the goals of the eu water framework directives, losses of n must be reduced, especially to coastal areas. this is true for the case study presented in this paper where mitigation actions are needed to considerably reduce the n impact on skive fjord, denmark, according to the recently published water plans for 2021–2027 (danish environmental protection agency 2023). nitrate (no3 –) in groundwater, surface water, the riparian zone and the sea floor can be reduced or retained (lutz et al. 2020). especially the longterm capacity of riparian zones in regulating groundwater nitrate fluxes is https://doi.org/10.34194/geusb.v53.8351 https://orcid.org/0000-0003-0982-4447 https://orcid.org/0000-0002-3796-3172 https://orcid.org/0000-0003-1882-2961 https://orcid.org/0000-0003-4094-4017 https://orcid.org/0000-0003-2318-145x mailto:psa@geus.dk sandersen et al. 2023: geus bulletin 53. 8351. https://doi.org/10.34194/geusb.v53.8351 2 of 14 www.geusbul let in.org not well understood (hill et al. 2014). nitrate removal in some riparian zones may be absent or minimal because deeper groundwater flow paths may not interact with nitrate-reducing minerals and organic matter (hill 2019). however, n flux to the marine environment might not be limited to output via streams and rivers. several studies suggest that transport of n to the marine environment via submarine groundwater discharge (sgd) can be substantial (e.g. bishop et  al. 2017; duque et al. 2019; taniguchi et al. 2019; szymczycha et  al. 2020), and its contribution to the local n budget may be much greater in geologically complex settings where fractures or preferential flow pathways can create hotspots or point sources of n fluxes (santos et al. 2021). a quantification of the amount of sgd under such geological settings, therefore, is challenging because in addition to direct measurements of groundwater flow, it requires detailed geological and hydrological models for both onshore and offshore (e.g. haider et al. 2015). likewise, assessments of the subsurface redox conditions are needed to determine where nitrate reduction takes place (hansen et  al. 2021; kim et al. 2021). therefore, modelling of geology, hydrology and geochemistry, along with calculations of the magnitude of the sgd and the amount of n transported to the coast is needed to evaluate the impact on the coastal environment. because the landscape and the near-surface geology of denmark are highly affected by the ice sheets and their meltwater during the pleistocene (e.g. houmark-nielsen 2011), the complex geological setting generally calls for advanced hydrogeological and redox modelling in order to map the groundwater flow paths and the fate of n (kim et al. 2019, 2021). buried tunnel valleys are structures that have a high impact on the geological architecture of the subsurface (jørgensen & sandersen 2006; sandersen & jørgensen 2017). most of these structures were formed as tunnel valleys eroded by high-pressure meltwater underneath the pleistocene glaciers. after formation, the valleys were filled with predominantly glacial tills and meltwater sediments (sandersen & jørgensen 2003). being up to 400 m deep and on average 1–1.5 km wide, buried valleys have been found to significantly impact the groundwater flow (e.g. sandersen & jørgensen 2003; andersen et al. 2013). several valleys are found close to the coastline (see sandersen & jørgensen 2022) and a few have been interpreted to cross the coastline (e.g. haider et al. 2015; andersen et al. 2016). in one example, it has been shown that groundwater in inferred buried onshore channels connects with shoreface sediments carrying nitrate to the coastal zone (andersen et al. 2007). in combination, these studies show a potential for buried valleys in enabling or enhancing preferential flow of groundwater across the coast, bypassing nitrate reduction zones in, for example, the riparian zones and thus transporting nitrate directly to the marine environment. in this paper, we present results from a multidisciplinary study on nitrate transport in the uppermost 40 m of the subsurface in a near-shore area on the salling peninsula, north-west denmark. here, large and deep as well as narrow and shallow buried tunnel valleys close to the shore are mapped in high detail. our study hypothesises that these mainly sand-filled and gravel-filled tunnel valleys play an important role as preferential pathways of nitrate to the coastal ecosystem, thereby increasing the risk of eutrophication. 2. site description 2.1. topography the study area comprises the two adjacent catchments of the streams hulebro bæk (study site 1; 11.3 km2) and hagens møllebæk (study site 2; 27.6 km2) located on the salling peninsula in the north-western part of jylland (fig. 1). the topography is highest to the north and in the south-west with elevations up to 40–50 m above sea level (a.s.l.), sloping down to below 20 m a.s.l. in the central and southern parts of the study area (fig. 1a). the two streams cut through the terrain in an overall south-easterly direction towards skive fjord (fig. 1) – a fjord that is generally 3 to 4 km deep, but up to 5 m deep in places and that constitutes a southern branch of the larger limfjorden farther north. limfjorden has connections to the kattegat to the east and the north sea to the west (fig. 1). 2.2. geological setting the pre-quaternary sedimentary succession relevant for this study consists of palaeogene clays with thin layers of miocene mica sand, overlain by silt (rasmussen et al. 2010). the overlying quaternary succession varies in thickness from less than 10 m to more than 120 m and is dominated by tills and meltwater deposits. the study area is characterised by a clay-dominated moraine landscape from the weichselian glaciation with  occurrences of postglacial freshwater deposits in low-lying areas. several ice advances have previously covered the area, resulting in a complex quaternary succession (houmark-nielsen 2011). the valley system previously mapped in the study area comprises two valleys with a n–sand nne– ssw-orientation and an older generation of valleys with a wnw–ese-orientation (fig. 2a). the valleys are mainly eroded into palaeogene clay-dominated sediments (down to 100 to 150 m below sea level [b.s.l.]). according to borehole data, the buried valleys are https://doi.org/10.34194/geusb.v53.8351 http://www.geusbulletin.org sandersen et al. 2023: geus bulletin 53. 8351. https://doi.org/10.34194/geusb.v53.8351 3 of 14 www.geusbul let in.org filled with successions of quaternary clay tills, meltwater sand and meltwater clays. all valleys are now completely buried and not directly recognisable in the present-day landscape (sandersen & jørgensen 2016b). 2.3. climate and agriculture the agricultural area in denmark spans about 61% (2022) of the country’s total area of 42 951 km2. the climate is costal temperate with a mean temperature of 9.1°c, and an average precipitation of 781.7 mm per year as reported by the danish meteorological institute for the years 2011–2020 (dmi 2023). the study area is dominated by agricultural fields (c. 88%), whilst buildings (c.  10%) and nature (c. 2%) occupy smaller parts of the area (levin et al. 2014). both sandy and clay-rich soil types are found within the study area (adhikari et al. 2013) with a relatively low content of organic matter (<6%; adhikari et  al. 2014). the crop types are mainly cereals and to a lesser extent permanent grass, and dairy, pig and arable farming types are abundant (danish agricultural agency 2018). 3. data and methods 3.1. drillings and lithological descriptions in this study, a drilling campaign operated by the ejlskov company was carried out to collect lithology and water chemistry samples at seven locations in november 2021. the lithology samples were collected using a geoprobe direct push method (dt-22), and the groundwater samples were collected using a modified geoprobe fig. 1 overview of the study area catchments. the northern catchment of hulebro bæk (study site 1, 11.3 km2) and the southern catchment of hagens møllebæk (study site 2, 27.6 km2) are shown. a: elevation in m a.s.l. b: surface geology (from jakobsen et al. 2022). c: location of geochemical sampling points (p1, p1a, p4, p4a, p5, p6 and p7). d: areas covered by the ttem soundings survey. legend surface geology (1:25 000) freshwater sand meltwater sand clayey till existing wells (jupiter) p6p6p7p7 p5p5 p1ap1a p1p1 p4p4 p4ap4a (b) (c) study sites 0 1 2 (a) hagens møllebæk hulebro bækn skive fjord lyby streamline limfjord km freshwater clay freshwater gyttja freshwater peat saltwater sand meltwater gravel meltwater clay gravelly till sandy till saltwater clay urban area no information oligocene/miocene/mica sand stream stations investigation wells (this study) ttem soundings 0 m (a.s.l) 56 m (a.s.l) kattegat north sea (d) https://doi.org/10.34194/geusb.v53.8351 http://www.geusbulletin.org sandersen et al. 2023: geus bulletin 53. 8351. https://doi.org/10.34194/geusb.v53.8351 4 of 14 www.geusbul let in.org fig. 2 geophysical data and geological interpretations. (a) buried valleys mapped prior to the present study (sandersen & jørgensen 2016a, b, 2017). (b) buried valleys mapped using the ttem method. interpreted from ttem 3d grid slices, with mean resistivity slices at 2 m a.s.l. (see resistivity scale in figure legend). (c) a south–north-orientated cross-section through the 3d ttem resistivity grid superimposed with geological interpretations from boreholes. bv: buried valley. meltwater sand/gravel well (no lith. info.) meltwater clay clay, sand, gravel clay till borehole lithology resistivity (ohm-metre) 1 10 100 1000ttem resistivity (3d grid) (map ’b’ and profile ’c’) water level screen 5. 12 3 silt ttem doi (depth of investigation; lower) (profile ’c’) 1 km 1 km n n centre line buried valley; well documented in data buried valley; poorly documented in data buried valley outline (valley name refers to profile ’c’)bv-a bv-a bv-b bv-c bv-d bv-e profile ’c’ study site 1 study site 2 roads borehole (national well archive at geus) borehole no. lill scale: 15.0 2.000 4.000 8.0006.000 0 50 –50 0 10.000 distance (m) e le va tio n (m a .s .l. ) 55 .2 81 6 55 .4 60 55 .4 38 46 .3 33 46 .1 30 7 46 .2 00 1 46 .2 00 0 46 .6 84 46 .7 02 46 .5 94 46 .8 16 46 .8 14 46 .8 15 46 .6 32 46 .1 88 46 .3 64 46 .5 29 south bv-a bv-d bv-ebv-b bv-c interpreted buried valleys bv-d north meltwater clay/till palaeogene clay meltwater sand/gravel meltwater sandclay sand clay sand/ gravel sand palaeogene clay clay/till clay clay sand clay c a b skive fjord skive fjord profile (c) profile (c) https://doi.org/10.34194/geusb.v53.8351 http://www.geusbulletin.org sandersen et al. 2023: geus bulletin 53. 8351. https://doi.org/10.34194/geusb.v53.8351 5 of 14 www.geusbul let in.org sp16 groundwater sampler (shorter screen, i.e. 0.25 m) and a peristaltic pump. older borehole data have been extracted from the freely accessible national borehole archive jupiter hosted by geus (geus 2022; hansen & pjetursson 2011). 3.2. geophysical surveys two types of geophysical tem (transient electromagnetic) data were used in this study: (1) ttem (towed tem) soundings data (auken et  al. 2019; aarhus university 2021) with a line spacing of 25 to 60 m shown in black on fig. 1d, and (2) tem40 (ground-based tem) soundings data (christiansen et al. 2006). the tem40 data are not included in fig. 1d. the tem method uses a transmitter coil to emit a current, which gives rise to a magnetic field. the current is then rapidly turned off, resulting in a decrease of the magnetic field, which, in turn, gives rise to a secondary magnetic field decaying over time. the rate of decay is measured in a receiver coil, which provides the resistivity of the subsurface. clay deposits are characterised by slow decays due to the conductive nature of the material, whereas sand or gravel layers, which are characterised by a high resistivity, will have fast decays. the two systems have different spatial resolutions and depths of investigation: ttem provides high-resolution images of the uppermost 70 m (auken et  al. 2019), whereas the depths of investigation for the tem40 system range from 200 to 300 m. as a result of this and in combination with the generally denser data coverage, highly detailed mapping of the near-surface geology is possible (sandersen et al. 2021). all tem systems are based on the time-domain electromagnetic method and measurement of the conductivity of the subsurface layers (danielsen et al. 2003). the method therefore provides information of the resistivity distribution of the sedimentary succession (sørensen & auken 2004). 3.3. stream discharge and water chemistry the stream discharge and nitrate concentrations were extracted from the public database oda (https://odaforalle.au.dk). for streams, water chemistry monitoring focuses only on nutrient compounds, such as nitrate, nitrite and phosphate. in both catchments, discharge and stream chemistry were monitored at the outlet of the catchment (fig. 1a). in hulebro bæk, stream discharge has been measured daily since 1989, and nitrate concentration has been monitored at monthly (dry season) to bi-weekly (wet season) intervals since 1984. in hagens møllebæk, discharge and stream water chemistry were monitored since 2020 (daily) and 2017 (bi-weekly to monthly), respectively. 3.4. groundwater chemistry the groundwater chemistry data were extracted from the jupiter database. in this study, the focus is on major cations (ca2+, mg2+ and na+), anions (cl–, no3 – and so4 2–), ph and the groundwater alkalinity. groundwater chemistry information was available from 48 boreholes (fig.  1c) within a depth range from 3 to 157 m b.l.s. (below land surface). the groundwater chemistry in these boreholes has been monitored periodically since 1951, and the data availability varied from 1 to 5 measurements per borehole. this study includes all the available measurements. except for nitrate, concentrations of major cations and anions are expressed in mm. stoichiometric ratios of these elements are used to infer underlying processes. nitrate, however, is shown in mg/l because the standard for groundwater and drinking water is 50 mg/l of nitrate according to danish and european legislation (vs. 0.8 mm), and nitrate concentrations below 1 mg/l (vs. 0.016 mm) are considered reduced conditions. thus, it is more intuitive to express nitrate in mg/l. 3.5. visualisation and interpretation of data the ttem soundings are visualised with low resistivities being blue to green and high resistivities being red to purple (see fig. 2). the doi (depth of investigation) shown in fig. 2c illustrates the maximum depth at which the ttem models can be considered reliable (christiansen & auken 2012). based on the ttem data, a 3d grid and a series of horizontal 2 m mean resistivity slices were made. the map in fig. 2b shows a selected resistivity slice at 2 m a.s.l., and fig. 2c shows a selected vertical cross-section through the 3d resistivity grid and selected boreholes. the tem40 data were used as single soundings to support the geological interpretations of the ttem data and are not visualised on the profile. the borehole data and the geophysical data were imported into the 3d geological modelling tool geoscene3d, enabling joint interpretations of the geological and geophysical data within the 3d model space (see figs 2b and c). 4. results 4.1. geology 4.1.1. the pre-quaternary succession the ttem mapping has confirmed the widespread presence of low resistive sediments with resistivities below 10–12 ohm-metre, in boreholes described as palaeogene clay. on top of this clay, sporadic sandy and clayey neogene sediments are found. the elevation of the top of the pre-quaternary sediments varies from 20 m a.s.l. https://doi.org/10.34194/geusb.v53.8351 http://www.geusbulletin.org https://odaforalle.au.dk https://odaforalle.au.dk sandersen et al. 2023: geus bulletin 53. 8351. https://doi.org/10.34194/geusb.v53.8351 6 of 14 www.geusbul let in.org to deeper than 100 m b.s.l. (fig. 2c), mainly due to erosion during the quaternary. 4.1.2. the quaternary succession and buried valleys the ttem survey has added more detail to the uppermost 50 to 70 m of the subsurface and has especially enhanced the picture of the buried valleys in the area (sandersen & kallesøe 2021). the deep and broad buried valleys that were earlier mapped primarily with tem40 and boreholes (see fig. 2a) were confirmed with the ttem survey. in addition to these valleys, several narrow and shallower valleys are mapped with ttem (figs 2b and c). a large n–s-orientated buried valley is seen on the horizontal mean resistivity slice 2 m a.s.l. in fig. 2b (labelled bv-e) as generally high resistivities (mainly red to yellow colours) surrounded by low resistivities (blue colours). the valley is clearly seen to the right on the cross-section in fig. 2c, where sandy and clayey quaternary sediments represent the infill of the c. 2 km wide structure eroded to depths below 100 m b.s.l. the valley is surrounded by the palaeogene clays showing very low resistivities. the presence of the valley below the ttem doi is confirmed by tem40 and boreholes. the data show that the deep parts of the valley (below c. 10 m b.s.l.) are divided into two parts, one dominated by clay and the other by sand, represented by low and high resistivities, respectively (fig. 2c). according to borehole data, the sediments consist of meltwater sands and clays. between c. 10 m b.s.l. and c. 30 m a.s.l., the ttem shows a 10–20 m thick low-resistive layer. the clay, generally interpreted as meltwater clay in borehole data, has a wavy appearance and is apparently confined to the valley. above the clay, there is a succession of meltwater sand or gravel irregularly covered with clayey till. according to previous investigations, the buried valley structure continues southwards to the coast of skive fjord (fig. 2a). a wnw–ese-orientated buried valley in the central part of the study area is also around 2 km wide eroded no deeper than c. 60 m b.s.l. (labelled bv-c in fig. 2). the infill is dominated by high resistivities representing meltwater sand and gravel. above sea level, clay layers are found occasionally. along sections of the northern valley flank, a 500–700 m wide high-resistivity structure can be observed from the terrain surface and down to c. 10 m b.s.l. the structure is split into a nw–se segment and a wsw–ene segment (bv-d in fig. 2b). at elevations higher than the resistivity slice shown in fig. 2b, another narrow high-resistivity structure is present above and with the same orientation as the deep, broad valley. both valleys are labelled bv-d on the cross-section in fig. 2c. according to previous investigations, the buried valley structure continues eastwards to the coast of skive fjord (fig. 2a). to the south, two narrow high-resistive structures with an overall wsw–ene orientation can be seen (fig.  2b). the northern of the two, bv-b, is 200–400 m wide, eroded down to 20 m b.s.l. and generally filled with high resistivity layers corresponding to sand. the deepest parts of the valley are eroded down to 50–60 m b.s.l. into the low-resistive palaeogene clays. the southern valley, bv-a (fig. 2b and c), appears to be two to three separate erosions, just 200 to 300 m wide, which merge eastwards to create a single valley trace. the ttem soundings map the southernmost valleys (bv-a) all the way to the coastline of skive fjord, whereas the valley bv-b, just to the north, is mapped as close as 800 m from the coast. most of the buried valleys in denmark are formed as buried valleys eroded by meltwater flow underneath the pleistocene ice sheets (jørgensen & sandersen 2006). the valleys mapped in this study area are no exception. the ttem mapping has revealed an additional system of younger, narrower and shallower buried valleys in the upper part of the subsurface (fig. 2b; valleys bv-a, bv-b and bv-d). as seen on the cross-section in fig. 2c, these valleys are filled with high-resistive sediments dominated by meltwater sand and gravel, which are only sporadically covered by low-resistive clayey sediments. all of the buried valleys in the area are several kilometres long and have been mapped very close to the shore of skive fjord to the east. 4.2. nitrate concentration–discharge relationships the nitrate concentration–discharge (c–q) relationships of both catchments show positive slopes up to a specific discharge of 1 mm/day and then negative slopes at higher discharge, likely due to dilution at extreme rainfall events (fig. 3). however, the nitrate concentrations at baseflow were much higher in hulebro bæk (c. 2.0 mg/l) than in hagens møllebæk (c. 5 mg/l). in addition, the positive slope of the c–q relationship was much steeper in hulebro bæk compared to that of hagens møllebæk (fig. 3). at hulebro bæk, the long-time series of nitrate data (1984–2020) reveals a clear reduction in stream nitrate concentrations, likely due to n regulation of agriculture since the 1980s. 4.3. groundwater chemistry the boreholes holding groundwater chemistry data were divided into groups representing the individual buried tunnel valleys (fig. 4). in all the valleys, groundwater is reduced (i.e. nitrate concentration <1 mg/l) at https://doi.org/10.34194/geusb.v53.8351 http://www.geusbulletin.org sandersen et al. 2023: geus bulletin 53. 8351. https://doi.org/10.34194/geusb.v53.8351 7 of 14 www.geusbul let in.org around 30 m b.l.s., roughly corresponding to sea level, except for groundwater in bv-c. at shallow depths (<30 m b.l.s.), nitrate concentrations in the bv-e valley showed the largest variability (0–100 mg/l) compared to concentrations elsewhere (fig. 4). groundwater nitrate in bv-c and bv-a increased with increasing depth (fig. 4). for instance, nitrate concentrations in groundwater from valley bv-c displayed 120 mg/l at around 30 m b.l.s. or at sea level, which is the highest concentration amongst all the groundwater chemistry measurements. groundwater from valley bv-d showed moderately high nitrate concentrations, which decreased with increasing depth. the concentrations of ca2+ + mg2+ in all groundwater samples varied between 2 and 4 mm, and those of hco3 – were 0.2–5 mm (fig. 5). the so4 2– concentrations ranged between 0.1 and 1.8 mm (fig. 5). those of na+ and cl– were 0.6–3.2 mm and 0.7–2.4 mm, respectively. groundwater ph varied between 7 and 8 (fig. 5), except for one case in valley bv-c (ph = 5.77, not shown in the figure). fig. 4 nitrate concentration in groundwater. on the map (a), the average of all existing data of each borehole is shown in colour. nitrate concentrations of groundwater are shown as a function of sampling depth (b) and elevation (c). 0 50 100 nitrate (mg/l) –100 –90 –80 –70 –60 –50 –40 –30 –20 –10 0 b el ow la nd s ur fa ce (m ) 0 50 100 nitrate (mg/l) –70 –60 –50 –40 –30 –20 –10 0 10 20 30 e le va tio n (m ) sea level 1 km n skive fiord bv-a bv-b bv-c bv-d bv-e 0 20 40 60 80 100 120 nitrate (mg/l) bv-a bv-b bv-c bv-d bv-e a b c fig. 3 nitrate concentration–discharge (c–q) relationships of the study catchments. the sampling year of the stream chemistry is shown in colour. 10–3 10–2 10–1 100 101 mm/day 1 10 25 50 100 150 ni tra te (m g/ l) 10–3 10–2 10–1 100 101 mm/day 1 10 25 50 100 150 ni tra te (m g/ l) hulebro bæk hagens møllebæk year 1990 2000 2010 2020 https://doi.org/10.34194/geusb.v53.8351 http://www.geusbulletin.org sandersen et al. 2023: geus bulletin 53. 8351. https://doi.org/10.34194/geusb.v53.8351 8 of 14 www.geusbul let in.org 5. discussion 5.1. nitrate pathways from land to coast via streams the c–q relationships provide valuable information about stream water and solute transport at the catchment scale (godsey et  al. 2009; kim et  al. 2017, 2021; bieroza et  al. 2018). the differing c–q relationships in the two catchments indicate that nitrate is transported via different pathways from the agriculture fields to the stream outlets. the lower nitrate concentration at baseflow and the steeper slope of c–q in hagens møllebæk compared to those of hulebro bæk can be interpreted as a limited contribution of groundwater to outlet (fig. 3). during baseflow, the stream hagens møllebæk may be fed by the groundwater in the riparian zone; thus, it likely contains very low nitrate. the steep slope of the c–q relationships of hagens møllebæk implies that nitrate is primarily transported via shallow subsurface pathways such as drains, interflow or overland flow in this system. the surface geology of the hagens møllebæk catchment, which is predominantly clayey till, supports this interpretation (fig. 1b). in hulebro bæk, groundwater appears to be significant in transporting nitrate to the stream. valley bv-c intersects with the hulebro bæk outlet (fig. 2), and the groundwater in this buried valley is oxic with nitrate concentrations exceeding 100 mg/l (fig. 4). this oxic, nitrate-containing groundwater may discharge into the stream year-round, resulting in higher nitrate concentration during baseflow and also the gentle slope of the c–q relationship shown in fig. 3. in addition, this catchment is covered with sandy till, and the subsurface contains meltwater sand and gravel, which explains the deeper percolation of nitrate containing groundwater directly discharging to the stream (figs 1b and 2). thus, in this catchment, shallow subsurface flow appears to play a minor role in delivering nitrate to the stream. 5.2. nitrate pathways from land to coast bypassing the stream outlets: the role of buried tunnel valleys the groundwater and stream chemistry observations of hulebro bæk clearly demonstrate an important role of groundwater for nitrate transport within the catchment. this implies that groundwater may be a direct pathway fig. 5 major cation and anion concentration of groundwater in buried tunnel valleys. valleys bv-a, bv-c, bv-d and bv-e are shown as triangles, circles, diamond, and squares, respectively. the colours of the markers represent nitrate concentrations. 0 2 4 6 ca2+ + mg2+ (mm) 0 1 2 3 4 5 6 h c o 3 ( m m ) – 0 2 4 60 1 2 3 4 5 6 s o 4 (m m ) 2– 0 2 4 6 na+ (mm) 0 1 2 3 4 5 6 6 7 8 9 ph 0 1 2 3 4 5 6 h c o 3 ( m m ) 0 20 40 60 80 100 120 nitrate (mg/l) sea sa lt 1:1 1: 2 1:1 2:1 bv-a bv-c bv-d bv-e ca2+ + mg2+ (mm) c l (m m ) – – https://doi.org/10.34194/geusb.v53.8351 http://www.geusbulletin.org sandersen et al. 2023: geus bulletin 53. 8351. https://doi.org/10.34194/geusb.v53.8351 9 of 14 www.geusbul let in.org to the near-coast waters, bypassing the stream as well. along valley bv-c, for instance, groundwater containing up to 120 mg/l of nitrate was found far from any stream outlet and very near the coast, and this groundwater is present 2–10 m b.s.l. the biogeochemical evolution of nitrate discharging via groundwater to the sea floor is spatially complex and temporally dynamic (e.g. kroeger & charette 2008); however, various studies have documented that sgd should be part of the nitrate budget of the coastal ecosystems (andersen et al. 2007; luijendijk et al. 2020; brookfield et al. 2021). the spatial heterogeneity of sgd is one of the biggest challenges in quantifying its contribution to biogeochemistry of coastal ecosystems (kroeger & charette 2008; luijendijk et al. 2020). in this study, we found a high degree of spatial heterogeneity. in valleys bv-c and bv-d, groundwater chemistry showed a signature of pyrite oxidation by oxygen as in eq. (1), coupled with carbonate dissolution by sulfuric acid as in eq. (2). this results in so4 2–/(ca2+ + mg2+) of c. 0.5 and hco3 –/( ca2+ + mg2+) of c. 1 (fig. 6; appelo & postma 2005): fes o h o fe oh so h 15 4 7 2 2 42 2 2 3 4 2–( )+ + = + + + (1) ca mg co h xca x mg hco 1– x x1– 3 2 2 3 –( )+ = + +( ) + + + (2) the pyrite content, however, may be insufficient to reduce nitrate. these two buried valleys are primarily filled with sand without protective clay layers (fig. 2), and therefore, it is possible that the infill of these buried valleys is more weathered and thus depleted of reduced compounds, that is, pyrite or reactive organic matter, than those of the other buried valleys. in valley bv-e, which is the largest and deepest buried valley in the study catchment, the groundwater chemistry varies widely. the nitrate containing groundwater of bv-e shows similar signatures as those of valleys bv-c and bv-d, except for two cases of high nitrate concentrations (so4 2–/(ca2+ + mg2+) of c. 0.1), which probably represent young oxic groundwater with a relatively low sulphate content due to a decrease in atmospheric deposition and absence of pyrite oxidation (fig. 5). we conclude that the upper oxic part of valley bv-e does not contain reduced compounds reactive enough to remove all the oxygen and initiate nitrate reduction. however, many, mostly deeper, groundwater samples were nitrate-free in valley bv-e (fig. 4). this groundwater shows a high variability: so4 2–/(ca2+ + mg2+) ranges between 0.1 and 0.5, and hco3 –/(ca2+ + mg2+) ranges between 1 and 3 (fig. 6). complete pyrite oxidation by nitrate reduction is defined in eq. (3). followed by carbonate dissolution, this will result in so4 2–/ (ca2+ + mg2+) = 2 and hco3 –/( ca2+ + mg2+) = 1 (appelo & postma 2005): fes no h o n feooh so h 5 15 5 0.5 5 10 5 2 3 – 2 2 4 2 – + + = + + + + (3) these stoichiometric ratios were not observed in our catchments; thus, we conclude that pyrite oxidation may not be responsible for the nitrate-free conditions. instead, we attribute these conditions to the complex geological structure of valley bv-e. in bv-e, the clay layers protect the groundwater from nitrate contamination, but because they are deformed, sandy geological windows most likely result in localised hydraulic pathways resulting in complex redox structures. the nitrate-free groundwater in valley bv-e may be well protected by the clay layers, whilst the oxic groundwater in bv-e may be related to sandy geological windows. the increase of hco3–/(ca2+ + mg2+) above 1, which is positively correlated to na+/cl–, may support this hypothesis (fig. 6). we attribute this positive correlation to cation exchange between ca2+ in groundwater and na+ in the salt-saturated marine sediments. the sediments of the protected areas may not be in equilibrium with the fresh groundwater because they have not been exposed to influx of enough fresh groundwater. especially in areas where buried valleys are eroded into high-lying, clay-dominated substrata and filled with coarse-grained sediments, groundwater flow will predominantly take place within the buried valleys. several deep and shallow buried valleys are mapped within the study area. infill ranges from disturbed layers of heterogeneous sandy, gravelly and clayey layers to an apparently homogeneous infill of sand and gravel (fig. 2). because the palaeogene clays outside the valleys can be found very close to the terrain surface, the groundwater flow will be focused in the hydraulically transmissive infill of the buried valleys. to the east in the study area where the valleys are very close to the coast, the groundwater flow will be towards the coast within the buried valleys. the focused flow results in high local fluxes implying that the groundwater emanates through the seabed with relatively high velocities so that nitrate reduction in the seabed may not be fast enough to remove all nitrate before it enters the sea. if the valleys terminate at the coast, for instance, due to later erosion, groundwater discharge will take place along the shore. however, if the valleys continue underneath the seabed, sgd may take place at some point offshore, depending on the transmissivity of the seabed. due to a lack of offshore data, the extension of the buried valleys underneath skive fjord is unknown. however, the onshore geophysical mapping and the general knowledge of buried valleys in denmark make it highly likely that the valleys have continuations under the fjord. for sgd to occur, the valley aquifers need to be ‘leaky’. where, on the other hand, deep groundwater https://doi.org/10.34194/geusb.v53.8351 http://www.geusbulletin.org sandersen et al. 2023: geus bulletin 53. 8351. https://doi.org/10.34194/geusb.v53.8351 10 of 14 www.geusbul let in.org within the buried valley has a clayey seal that prevents salinisation, fresh groundwater within the valley infill could potentially form an overlooked resource of offshore groundwater (referred to as offshore fresh groundwater, ofg; micallef et al. 2020; zamrsky et al. 2022). fig. 6 ratios of hco3 – / (ca2+ + mg2+) vs. so4 2– / (ca2+ + mg2+) and na+ / cl– of groundwater in the buried tunnel valleys. nitrate concentrations of each groundwater sample are shown as colours. 0 1 2 3 0 0.5 1 0 1 2 3 0 0.5 1 0 1 2 3 0 0.5 1 0 1 2 3 0 0.5 1 0 1 2 3 0 0.5 1 1.5 2 0 1 2 3 0 0.5 1 1.5 2 0 1 2 3 0 0.5 1 1.5 2 0 1 2 3 0 0.5 1 1.5 2 sea salt c ar bo na te + s tro ng a ci d c ar bo na te + c o 2 pyrite + o2 + carbonate hco3 / (ca2+ + mg2+)– hco3 / (ca2+ + mg2+)– hco3 / (ca2+ + mg2+)– hco3 / (ca2+ + mg2+)– hco3 / (ca2+ + mg2+)– hco3 / (ca2+ + mg2+)– hco3 / (ca2+ + mg2+)– hco3 / (ca2+ + mg2+)– s o 4 / (c a2+ + m g2+ ) 2– s o 4 / (c a2+ + m g2+ ) 2– s o 4 / (c a2+ + m g2+ ) 2– s o 4 / (c a2+ + m g2+ ) 2– n a+ / c l– n a+ / c l– n a+ / c l– n a+ / c l– 0 20 40 60 80 100 120 nitrate (mg/l) bv-e bv-d bv-c bv-a https://doi.org/10.34194/geusb.v53.8351 http://www.geusbulletin.org sandersen et al. 2023: geus bulletin 53. 8351. https://doi.org/10.34194/geusb.v53.8351 11 of 14 www.geusbul let in.org 5.3. buried valley occurrence and impact on the subsurface architecture the map in fig. 7 shows the current knowledge of the location of buried valleys in denmark based on interpretations of tem data primarily in combination with borehole data (sandersen & jørgensen 2022). although only c. 50% of the onshore area is mapped with tem, fig. 7 shows many valleys close to the coast. based on comparisons between onshore and offshore tunnel valleys in the danish north sea, andersen et al. (2012) concluded that valley morphologies on both sides of the shoreline were comparable, and as today’s coastline does not constitute a natural limit for the extension of the buried valleys, it is highly likely that many of the danish onshore valleys extend offshore. buried valleys in denmark are infilled with mainly tills and meltwater sediments. rough calculations based on borehole lithologies show that around 40–75% of the infill is composed of meltwater sand and gravel, with the remaining part of the infill consisting of silty and clayey meltwater sediments and tills (sandersen & jørgensen 2003). because the tunnel valleys are eroded deeply into the subsurface and generally have heterogeneous infill, the impact on the architecture of the subsurface is generally significant (sandersen & jørgensen 2017). the coarse-grained meltwater sediments often constitute important, deep-seated aquifers, but due to the erosive nature of the valleys and the typically heterogenic infill, the groundwater resources are generally vulnerable to contaminants from the surface (sandersen & jørgensen 2003). 5.4. structural analogies and perspectives the buried valleys represent geological structures that are expected to expand beyond the coastline, and the aquifers contained within them therefore constitute a part of the present-day hydrological cycle. this implies a continuous flow of fresh groundwater and a flux of dissolved compounds to coastal areas through the valley sediments. potential flow of nitrate-containing groundwater directly to the coastline and the sea floor, by-passing the riparian zone, is currently not considered in danish nitrogen budgets in the water plans for the protection of the aquatic environment. nitrate is the major form of n in groundwater. however, it might be converted to reduced forms when discharging through the seabed depending on the balance between flow rates and nitrate reduction rates (andersen et al. 2007). measurement of the total n content in the bottom water fig. 7 buried valleys in the danish subsurface mapped 2021 (data from http://buriedvalleys.dk/; sandersen & jørgensen 2022). green lines mark the location and orientation of the mapped valleys. 9˚0´ 56˚0´ limfjord study area kattegat n or th s ea denmark germany n 50 km https://doi.org/10.34194/geusb.v53.8351 http://www.geusbulletin.org http://buriedvalleys.dk/ sandersen et al. 2023: geus bulletin 53. 8351. https://doi.org/10.34194/geusb.v53.8351 12 of 14 www.geusbul let in.org column below the halocline of the inner fjord has shown that this could be a considerable source of n for biological production creating oxygen deficiency in the summer period (gertz et  al. 2018). direct contributions of nitrate to the near-coastal environment via groundwater could be significant and might therefore be important for the choice of the right mitigation measures in agricultural areas to reduce the n impact on the marine environment. buried tunnel valleys are common both onshore and offshore in former glaciated areas (van der vegt et al. 2012). since many are found close to the coastline, their inferred continuations are obvious targets for both sgd and ofg investigations. other geological structures, which in equal ways may enable or enhance groundwater flow across the coastline, could be fault zones, graben structures and offshore continuations of large aquifers in porous rocks and sediments (e.g. edmunds et al. 2001; hinsby et al. 2001; cantarero et al. 2019; santos et al. 2021). outside denmark, the number of investiga tions of both sgd and ofg is increasing (see taniguchi et al. 2019; micallef et al. 2020). pertaining to the conditions in denmark, widespread near-surface chalk and limestone aquifers could also be potential targets for further investigations. in relation to ofg, the buried tunnel valleys potentially represent a renewable groundwater resource that has not previously been considered. this contrasts with the potential ofg resources described by zamrsky et  al. (2022) who regard most fresh offshore groundwater as likely non-renewable resources. this is because, at a global scale, they are mostly located far from the shore and thus interpreted as being detached from inland aquifers. 6. conclusions the near-surface geology in the study area is dominated by buried tunnel valleys. new detailed geological mapping has revealed near-surface valleys that are narrower and less deeply eroded in comparison to the previously known valleys. all buried valleys are found very close to the shore of skive fjord, indicating that the valleys most likely extend offshore. groundwater and stream water chemistry observations in the study area clearly demonstrate the important role of groundwater in nitrate transport within the catchment, and that groundwater in buried valleys may constitute a direct pathway to the coast, bypassing the stream and riverine systems. groundwater samples from the valley infills revealed that nitrate is present below the sea level within the buried valley infill. samples with as high as 50–120 mg/l of nitrate were found far from the stream outlet down to 10 m b.s.l. close to the coastline. the complex geological structure of the buried tunnel valleys appears to be responsible for the spatial heterogeneity of the groundwater geochemistry in this area, where sandy geological windows result in localised hydraulic pathways, resulting in complex redox structures. in addition to this, the buried valleys generally control the groundwater flow towards the coast. consequently, there is a high risk that n will discharge to the coastal waters of limfjorden via buried valleys. as such, this flux will represent a n contribution to the environment that is usually not included in estimates of n transported to the inner danish marine environment. based on our findings, we therefore suggest that direct groundwater discharge to the coast is included in the n management plans in areas such as these and in comparable geological settings. we also stress the importance of producing subsurface models, which include detailed delineation of geology and geochemistry. otherwise, calculations of the flux of n to the marine environment will not encompass all sources. the number of investigations of both sgd and ofg is rising internationally, and with the large number of buried tunnel valleys found in denmark, there appears to be a potential for further studies, here. to calculate and evaluate actual n fluxes to the shallow marine environment, joint investigations should include detailed onshore and offshore mapping of buried tunnel valleys, groundwater modelling and geochemical investigations of both groundwater and sea water. apart from the buried tunnel valleys in former glaciated areas, fault zones, graben structures and offshore continuations of porous sediments with high flow rates and presumably low nitrate reduction rates would also be potential targets for further investigations. acknowledgements we wish to acknowledge the mapfield project (field scale mapping for targeted n-regulation and management), funded by the danish innovation fund (award number 8855-00025), for allowing us to publish project results. we also acknowledge the geological survey of denmark and greenland (geus) for supporting the publication of the paper. we are grateful for the comments and suggestions provided by the two reviewers. their suggestions greatly improved the manuscript. additional information funding statement the mapfield project ‘field scale mapping for targeted n-regulation and management’ was funded by the danish innovation fund (award number 8855-00025). author contributions ps, hk and bh conceptualised and designed the study. hk, rj and bh performed the fieldwork and interpretations related to geochemistry. jp performed fieldwork and data processing related to the geophysical https://doi.org/10.34194/geusb.v53.8351 http://www.geusbulletin.org sandersen et al. 2023: geus bulletin 53. 8351. 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(a) buried valleys mapped prior to the prese fig. 3 nitrate concentration-discharge (c-q) relationships of the study catchments. the sampling yea fig. 4 nitrate concentration in groundwater. on the map (a), the average of all existing data of eac fig. 5 major cation and anion concentration of groundwater in buried tunnel valleys. valleys bv-a, b fig. 6 ratios of hco3/ (ca2+ + mg2+) vs. so42/ (ca2+ + mg2+) and na+ / clof groundwater in the fig. 7 buried valleys in the danish subsurface mapped 2021 (data from http://buriedvalleys.dk/; sand data article | short koopmans et al. 2021: geus bulletin 47. 6526. https://doi.org/10.34194/geusb.v47.6526 1 of 7 new insights from field observations of the younger giant dyke complex and mafic lamprophyres of the gardar province on tuttutooq island, south greenland lot koopmans*1 , robert a. webster1 , rory changleng1 , lucy mathieson1 , alasdair j. murphy1 , adrian a. finch1 , william mccarthy1 1school of earth and environmental sciences, university of st andrews, st andrews, united kingdom abstract the gardar province of south greenland is defined by the products of alkaline igneous magmatism during the mesoproterozoic. the most laterally extensive gardar intrusions are a series of giant dyke complexes best exposed on the tuttutooq archipelago. we present new field observations and a geological map of north-east tuttutooq island that provide fresh insights into the temporal evolution of the younger giant dyke complex and two associated ultramafic lamprophyres. our data demonstrate that distinctive crystallisation regimes occurred in different sectors of the dyke complex, leading to the formation of marginal gabbros and ovoid pod-like domains displaying lamination, modal layering and/or more evolved differentiates. we infer that at least two pulses of magma contributed to the formation of the younger giant dyke complex. in addition, the relative ages of two ultramafic lamprophyre diatremes are constrained and attributed to two distinct phases of rifting in the gardar province. introduction determining how magmas intrude and crystallise in the crust using structural data and textural features provides fundamental insights into the processes inherent in the construction of igneous intrusions. despite decades of research, the mechanisms that control layer formation, magma transport and storage, and in-situ magma evolution in magma chambers remain elusive (e.g. cashman & giordano 2014). the younger giant dyke complex (ygdc) is an elongate igneous intrusion that extends for over 145 km across southern greenland (upton 2013). the ygdc contains regions in which igneous layering, foliation and/or differentiated compositions are observed (upton 1962). it therefore provides a unique opportunity to study these enigmatic magmatic processes. this study presents a detailed field analysis of eastern tuttutooq island, focusing on the sections of the ygdc exposed on the island and two ultramafic lamprophyre intrusions. we use these data to elucidate the structural and crystallisation history of the ygdc magma and associated lamprophyres, and thereby gain insights into the processes that controlled the formation of these spectacular intrusions. *correspondence: lk55@st-andrews.ac.uk received: 05 feb 2021 accepted: 20 apr 2021 published: 16 june 2021 keywords: gardar province, rift magmatism, younger giant dyke complex, magma chamber, lamprophyre abbreviations: bfd: big feldspar dykes ogdc: older giant dyke complex tin: tuttutooq-ilimmaasaq-narsarsuaq ygdc: younger giant dyke complex geus bulletin is an open access, peerreviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: kristian svennevig (geus, denmark) reviewed by: lotte melchior larsen (geus, denmark), and one anonymous reviewer funding: see page 7 competing interests: none declared additional files: see page 7 https://doi.org/10.34194/geusb.v47.6526 https://orcid.org/0000-0002-4018-2588 https://orcid.org/0000-0002-9149-6106 https://orcid.org/0000-0001-7660-1927 https://orcid.org/0000-0002-8840-7926 https://orcid.org/0000-0002-0920-1912 https://orcid.org/0000-0002-3689-1517 https://orcid.org/0000-0002-7214-1449 mailto:lk55@st-andrews.ac.uk koopmans et al. 2021: geus bulletin 47. 6526. https://doi.org/10.34194/geusb.v47.6526 2 of 7 www.geusbul let in.org geological background the tuttutooq–ilimmaasaq–narsarsuaq (tin) zone is the southern of two ene-orientated rift zones that comprise the gardar province. rifting occurred during the break-up of the columbia supercontinent, and magmatism was confined to two distinct periods: early (1320–1250 ma) and late (1200–1140 ma) gardar (upton 2013). today, the tin zone is composed of a granitic basement (julianehåb batholith), a fault-bound sedimentary and extrusive volcanic sequence (eriksfjord formation), and a range of intrusive bodies, of which those exposed on tuttutooq are presented in table 1. apart from minor occurrences of brown dykes, gardar magmatism on tuttutooq began with the emplacement of the older giant dyke complex (ogdc) – a composite giant dyke with syenogabbroic margins and a central zone grading from augite syenite in the west, to peralkaline sodalite-nepheline syenite in the east. this predates a suite of subvertical dykes, obtaining widths of up to 800 m, known as the ygdc. the ygdc is dominantly troctolitic but encloses several large ovoid pods of layered and/or more evolved rocks aligned parallel to the strike of the dyke (upton 1962; upton et al. 2003). following the giant dykes, the main gardar dyke swarm was emplaced throughout the tin zone (upton 2013). relatively small occurrences of ultramafic lamprophyre, occasionally associated with breccia pipes and carbonatites, are thought to have been emplaced throughout rifting. two of these are preserved on tuttutooq (upton et al. 2006). the latest known igneous events were the emplacement of large central complexes along the rift axis c.1160–1140 ma, including the tugtutôq central complex and ilímaussaq complex (waight et al. 2002). after rifting ceased, the gardar province remained remarkably well preserved. block faulting, erosion and uplift have set present day exposure at an estimated depth of 3–4 km below the contemporary land surface (upton 2013). it is thought that tuttutooq sits structurally below the nearby ilímaussaq complex (upton 2013), and hence the island provides an opportunity to study deeper rift plumbing systems below economically critical deposits. results a new geological map produced during the current work is presented in fig. 1. mapping was conducted at the 1:5000 scale, and a full resolution copy can be found in the supplementary files. in the following section, we provide a summary of field relationships within the ygdc and two ultramafic lamprophyre diatremes in the study area. younger giant dyke complex ygdc facies description the basis for subdivisions of lithological units used in this study is provided later. ygdc marginal gabbro. the ygdc marginal gabbro is a fineto medium-grained gabbro. dendritic plagioclase (<3 cm) and leucocratic melt segregate lenses are observed; the long axis of both are perpendicular to the contacts between host rock and ygdc at several localities (figs. 2b, d). rounded host rock xenoliths up to 15 cm are also observed sporadically within this facies (fig. 2a). ygdc central troctolite. a coarse-grained troctolite comprised of predominantly olivine and plagioclase primocrysts. minor oxides and interstitial clinopyroxenes are also present. four variations of the ygdc central troctolite exist: (1) a texturally homogeneous facies typically with <1% anorthosite xenoliths; (2) an anorthosite xenolith facies with 50–80% anorthosite xenoliths (0.1–150 m) surrounded by a troctolite matrix; (3) a glomerocrystic facies comprising a framework of 2–3 cm ‘snowflake’-like plagioclase aggregates; (4) a layered facies, which has modal layering and/or a foliation defined by a plagioclase crystal framework. ygdc syenogabbro. alkali feldspar-bearing gabbro with primocrysts of clinopyroxene. assorutit syenogabbro has characteristic dendritic clinopyroxene (<30 cm). krydssø syenogabbro grades from modally layered through foliated, to texturally homogeneous (l3, fig. 1). table 1 gardar intrusions on tuttutooq name rock unit trend age (ma) ± (ma) reference brown dykes (bd0) dolerite wnw–ese 1279 1.3 upton 2013 older giant dyke nepheline syenite e–w 1184 5 heaman (unpublished data) younger giant dyke troctolite ene–wsw 1163 2 heaman (unpublished data) gardar dyke swarm various, including big feldspar dykes (bfds) ene–wsw 1163–1160 upton 2013 (relative age) tuttutooq ultramafic lamprophyre mela-aillikite n/a <1163 upton et al. 2006 (relative age) tuttutooq central complex quartz syenites, granites n/a 1143 36 blaxland et al. 1978 see upton 2013 for original references and isotopic systems in detail. https://doi.org/10.34194/geusb.v47.6526 http://www.geusbulletin.org koopmans et al. 2021: geus bulletin 47. 6526. https://doi.org/10.34194/geusb.v47.6526 3 of 7 www.geusbul let in.org 400 200 20 0 200 300 20 0 200 20 0 100 100 10 0 20 0 200 10 0 20020 0 10 00 0 10 0 100 20 0 100 100 200 20 0 200 20 0 100 200 100 100 10 0 300 200 20 0 30 0 20 0 300 20 0 200 200 300 300 200 200 30 0 300 200 30 0 300 20 0 200 200 10 0 200 20 0 20 0 200 0 0 0 100 0 100 0 100 10 0 100 100 100 10 0 100 0 10 0 0 10 0 100 200 200 100 100 100 100 100 100 100 51 3524 30 86 18 20 38 57 5042 79 24 16 25 8 46°7'0"w46°8'0"w46°9'0"w46°10'0"w46°11'0"w46°12'0"w46°13'0"w46°14'0"w46°15'0"w46°16'0"w 60 °5 4' 0" n 60 °5 3' 0" n 60 °5 2' 0" n 0 1 2 km l4 assorutit l2 sissarluttooq layering and foliation contacts sharp fault gradational structural data horizontal layering julianehåb granite ygdc marginal gabbro late dyke swarm geological units late gardar units ygdc syenogabbro (homogeneous) ygdc alkali feldspar syenite anorthosite xenolith troctolite ygdc central troctolite (homogenous) glomerocrystic troctolite evolved veins assorutit syenogabbro (dendritic px) krydssø syenogabbro (foliated) krydssø syenogabbro (layered) layered troctolite younger giant dyke complex nepheline syenite syenogabbro older giant dyke complex early gardar units brown dykes lamprophyre localities basement assorutit sissarluttooq nasaasarli narsaq sund tunuaraqiteq marraat afersuaq l3 krydssø l1 marraat not mapped kryd ss ø store pilesø fig. 2a fig. 3g fig. 3afig. 3e fig. 3d fig. 2d fig. 2b fig. 2c fig. 3b fig. 3c fig. 1 composite map of the studied area, with location identified by a red arrow in the inset map. lamprophyre locality maps are presented in fig. 4. localities 1–4 (circled in red) refer to select evolved or layered ‘pods’ and are discussed in the text. a granite xenolith melt segregate lenses b c) marginal facies central facies c dendritic plagioclase central facies marginal facies d fig. 2 selected photographs of key exposure within the ygdc. a: granite xenolith incorporated in the ygdc marginal gabbro. b: melt segregate lenses in the marginal gabbro perpendicular to contact at sissarluttooq. c: dendritic plagioclases within the marginal gabbro sharply truncated against the central facies troctolite. d: wavy contact between marginal gabbro and central troctolite. the north arrow is 15 cm, and the hammer handle is 10 cm. location of each image can be found in fig. 1. ygdc alkali feldspar syenite. alkali feldspar-dominated syenite with clinopyroxene, quartz (<5%), amphibole, olivine, and calcite. evolved veins (commonly <1 m wide) are similar but contain more quartz (up to 20%) and are occasionally pegmatitic. ygdc facies relationships two sections (henceforth limbs) of the ygdc are present in the study area. both have subvertical, sharp contacts with the basement granite. they have wider (up to 800 m) and narrower sections (minimum 150 m), and trend ene–wsw. https://doi.org/10.34194/geusb.v47.6526 http://www.geusbulletin.org koopmans et al. 2021: geus bulletin 47. 6526. https://doi.org/10.34194/geusb.v47.6526 4 of 7 www.geusbul let in.org the outermost 2–100 m of each limb consists of the ygdc marginal gabbro, which shares a lobate or diffuse contact with the homogeneous facies troctolite (figs. 2c, d). the gross morphologies of the limbs are therefore equivalent. several petrographically distinct ovoid pods occur in the centre of the widest parts of the ygdc (fig. 1). contacts between the homogeneous facies troctolite and the pods are not exposed. where pods are concentrically zoned, internal contacts are gradational over c. 3 m. four individual pods are discussed in detail later. marraat pod (l1) the marraat pod is 300 m wide, is composed of layered troctolite and in places sheathed with glomerocrystic troctolite (l1, fig. 1). contacts between the two facies are complex, and occasional ‘glomerocrysts’ are found within layers. within the layered troctolite, alternating picritic and thicker troctolitic layers dip towards the centre of the pod, defining a synformal structure. layers are laterally continuous and typically thicker at the axis (up to 30 cm) than the edges (1 cm). contacts between layers are generally gradational over 1  cm; however, some sharp boundaries are observed to the south of the pod where the layers steepen to near vertical. feldspar crystals define a weak layer-parallel foliation within the troctolite and a non-pervasive, layer-parallel foliation in the picrite that is most apparent at the southern margin. pegmatitic autoliths of the ygdc central troctolite (up to 30 cm) occur in the southern section of the layered pod. sissarluttooq pod (l2) the sissarluttooq pod is dominated by a layered troctolite that is best exposed along the eastern coast of tuttutooq, where the ygdc attains a width of 800 m (l2, fig. 1). a well-developed foliation is persistent throughout the pod defined by a framework of plagioclase. foliations dip concentrically inwards and shallow from 50° at the edge towards a horizontal central point, defining a canoe-shaped morphology. crystal size is uniform throughout the pod. modal layering is observed at only one locality where decimetre-scale picritic layers alternate with foliated troctolite across a 5 m interval. elongate mafic enclaves composed of olivine-magnetite cumulates (60–40%) with interstitial plagioclase are observed both aligned with (up to 6 m long) and cross-cutting the foliation (0.15–1 m long). magnetite crystals are often aggregated and have poikilitic textures enclosing rounded olivine and/or plagioclase crystals. enclave margins are cuspate against the troctolite, are not chilled (fig. 3), and entrain ellipsoidal autoliths or individual phenocrysts from the neighbouring-foliated troctolite. a thin (4–50 m) sheet of homogeneous facies troctolite dissects the ygdc through the sissarluttooq pod (fig. 3a). it has a fine-grained chilled margin and terminates bluntly outside of the pod (fig. 3a). a b c d e f g homogenous finger foliated troctolite chilled margin north foliated troctolite enclaves enclave foliated troctolite anorthosite evolved vein kryddsø syenogabbro north layered troctolite fig. 3 a: contact between a ‘homogenous finger’ and foliated troctolite at sissarluttooq. b: modal layering in the krydssø pod. c: overview of layered synform in the marraat pod. d: foliated troctolite and parallel mafic enclaves in the sissarluttooq pod. e: large mafic enclave entraining portions of foliated troctolite in the sissarluttooq pod. f: bilateral veins (e.g. white arrow) within mafic lamprophyres. g: evolved vein cross-cutting anorthosite at assorutit. the north arrow is 15 cm, and the thumb is 2 cm. location of each image can be found in fig. 1. https://doi.org/10.34194/geusb.v47.6526 http://www.geusbulletin.org koopmans et al. 2021: geus bulletin 47. 6526. https://doi.org/10.34194/geusb.v47.6526 5 of 7 www.geusbul let in.org krydssø pod (l3) exposed c. 1 km ssw of assorutit (l3, fig. 1), this pod is composed of krydssø syenogabbro. thick modal layering (5–15 cm) of laterally continuous feldspar-rich (<80% mode) and feldspar-poor (>40% mode) layers is exhibited in the western third of the pod. to the east, layering grades into a foliation defined by euhedral aligned 2–10 mm pyroxene (c. 20% mode), before diminishing into texturally homogeneous krydssø syenogabbro. layering and foliation both dip from c. 40° at the margins towards the centre of the pod and define a central synformal axis plunging 10° towards the north-east. two large evolved veins (up to 60 m wide) have been identified near the krydssø pod. assorutit pod (l4) the assorutit pod (l4, fig. 1) grades concentrically from the homogeneous facies troctolite through 50–100 m of assorutit syenogabbro to a core of alkali feldspar syenite. towards the tip of assorutit peninsula, an anorthosite xenolith troctolite is juxtaposed against the pod by a wnw–ese fault marked by a 2–5 m wide zone of brecciated troctolite. pegmatitic (0.5–10 cm) evolved veins cross-cut and radiate outwards from the ygdc alkali feldspar syenite facies and can be observed on either side of the fault. the veins also cross-cut the large anorthosite xenoliths (fig. 3g). ultramafic lamprophyre two ultramafic lamprophyres are associated with the northern limb of the ygdc (fig. 4). eastern lamprophyre and breccia the eastern lamprophyre and an associated breccia diatreme outcrop at the apex of a narrowing in the ygdc, both are circular in map view with diameters of c. 100 m (fig. 4). the lamprophyre’s mineralogy consists of olivine, pyroxene, phlogopite and garnet, and occasional phlogopite pegmatitic segregations. cross-cutting, sinuous and bilateral veins (fig. 3) are observed radiating from the centre. the breccia pipe comprises clasts with rounded edges and a matrix that grades from a felsic to mafic composition. clasts include quartzite, a finegrained mafic lithology and coarse granite. the lamprophyre cross-cuts the breccia pipe causing further brecciation but no contact between either unit with the ygdc is exposed. dykes from the main gardar dyke swarm have variable relationships with the units. trachytic dykes cross-cut both the lamprophyre and breccia. however, a doleritic dyke with large felspar xenocrysts (i.e. a big feldspar dyke; bfd) is traceable for less than a metre into the breccia before it becomes indistinguishable from the surrounding breccia matrix (fig. 4). the matrix at this point is mafic in composition. cross-cutting relationships that nearby indicate the doleritic bfds were some of the earliest members of the main gardar dyke swarm, emplaced prior to the trachytic dykes. western lamprophyre the western lamprophyre is found wholly within the  ygdc near the northern contact with the ogdc. exposure is limited, but the unit appears to have an breccia (mafic host) trachyte dyke ultramafic lamprophyre breccia (felsic host) ygdc evolved veins big feldspar dyke, doleritic julianehåb granite ogdc syenogabbro ogdc nepheline syenite ygdc central troctolite contacts observed inferred gradational 46°11'25"w46°11'30"w46°11'35"w46°11'40"w 60 °5 4' 25 "n 46°14'50"w46°14'55"w46°15'0"w46°15'5"w 60 °5 3' 40 "n 60 °5 3' 35 "n 50 m 50 m western eastern fig. 4 focus maps of the western and eastern lamprophyre bodies (see fig. 1 for locations). https://doi.org/10.34194/geusb.v47.6526 http://www.geusbulletin.org koopmans et al. 2021: geus bulletin 47. 6526. https://doi.org/10.34194/geusb.v47.6526 6 of 7 www.geusbul let in.org irregular shape. the mineralogy can be divided into two groups: (1) oikocrysts consisting of 2–10 mm plagioclase, amphibole and biotite in the south and <2 mm biotite crystals in the north and (2) uniformly sized mafic chadacrysts that are enclosed by group 1. the northern contact with the ygdc is sharp and curviplanar, whereas the southern contact displays a complex melange of troctolite and lamprophyre. in one outcrop, orbicular minerals (potentially pseudo-leucite) are present. discussion origin of layering the layered pods of the ygdc have previously been interpreted as regions where convective overturn was vigorous enough to produce dynamic layering (upton 1962, 2013). however, we find the structural, mineralogical and textural differences between the marraat and sissarluttooq pods sufficient to warrant the presence of two different layer-forming processes. the synformal structure with layers thickening towards the centre at marraat may be explained by crystal-laden density currents originating from gravitational instabilities at the walls, implying convection may not be required. a similar process has been suggested to occur in parts of the skaergaard intrusion (irvine 1980), as well as other regions of the ygdc (upton 2013). this process is inconsistent with the uniform crystal size and concentric structure at sissarluttooq; instead, we hypothesise layering occurs here through gravitational settling below a convective cell. petrogenesis of ygdc magmas assorutit and krydssø magmas the presence of evolved rocks at the krydssø and assorutit pods has been attributed to either in-situ fractionation of the troctolitic melt (upton 1962) or the emplacement of distinct later injections of magma (upton & thomas 1980). field results presented here are not exclusive to either process. however, the layering mechanisms discussed earlier could fractionate a melt by mechanically separating crystals (i.e. dynamic fractional crystallisation). this process could occur during layer formation in the krydssø and sissarluttooq pods, driving the fractionation of the ygdc to more evolved compositions. furthermore, the presence of cross-cutting evolved veins within the anorthosite xenolith troctolite, known to be a part of the ygdc roof zone (upton 2013), indicates that silicic melt migrated within the ygdc along vein networks. this presents a process by which the central complexes of the tin zone could be fed by evolved melts fractionated at depth. the silica-saturated nature of the assorutit pod, in otherwise undersaturated lithologies, has been postulated to be the result of crustal assimilation (upton 2013). granite xenoliths found in this study provide direct evidence that this process was occurring. mafic enclaves we interpret the mafic enclaves at sissarluttooq to be cogenetic with the ygdc, and not later melt injections, due to the lack of chilled margins. such enclaves may be the result of the ygdc melt crossing into the fe-immiscibility field during fractional crystallisation. the resulting dense, fe-rich melts sank through the crystal mush below, whilst the lighter melts ponded at the top of the intrusion, forming the anorthosite. a similar process has been suggested to occur in the isortôq giant dykes (rosa et al. 2020). homogeneous sheet the presence of central facies troctolite that cross-cuts the sissarluttooq pod (l2) indicates that minor troctolitic injections occurred over a prolonged time period. both injections of ygdc melt are thought to have the same source due to their mineralogical and textural similarities. timing of lamprophyre magmatism the gradational change of a felsic to mafic matrix in the breccia diatreme at the eastern lamprophyre pipe is interpreted to record mixing between a primary felsic material and a mafic magma. we suggest that the felsic material formed through the melting of julianehåb granite and the mafic host was sourced from the bfd and possibly the lamprophyre. this requires the emplacement of these bodies before the breccia’s matrix had fully crystallised, which provides a tight relative age on the emplacement of the lamprophyre – contemporaneous with the emplacement of the bfds at the start of the main gardar dyke swarm. the equigranular ultramafic mineralogy present in the western lamprophyre implies a consistent crystallisation regime. in contrast, the variable size of the enclosing poikilitic mineralogy suggests that the minerals crystallised within a high thermal gradient. the textural differences between the north and south contacts of this lamprophyre support the presence of a high thermal gradient, which may be explained by the proximity to the external contact of the ygdc. in addition, the complex southern contact of the lamprophyre provides no evidence of a liquid–liquid contact between the lamprophyre and the ygdc. we suggest this body is a xenolith of an ultramafic lamprophyre emplaced just before the ygdc, producing an equigranular mafic mineralogy. density differences have led it to sink through the ygdc post-emplacement, generating the poikilitic mineralogy due to the interaction with the host magma. conclusion our field results lead to several conclusions: (1) density currents and/or convection developed the marraat and sissarluttooq pods; (2) the ygdc melts were available https://doi.org/10.34194/geusb.v47.6526 http://www.geusbulletin.org koopmans et al. 2021: geus bulletin 47. 6526. https://doi.org/10.34194/geusb.v47.6526 7 of 7 www.geusbul let in.org for a prolonged time period; (3) magmatic fractionation drove the formation of the assorutit and krydssø pods as a result of crystal settling under dynamic conditions; (4) the presence of evolved veins radiating from the assorutit pod demonstrates that there was mobilisation of magmatic differentiates within the ygdc; (5) the two ultramafic lamprophyres on tuttutooq are of two distinct ages in the late gardar (>1163 and 1163–1160 ma). targeted geochemical, anisotropy of magnetic susceptibility and spatial analyses will enhance our understanding of the ygdc petrogenesis presented here and further study the associated lamprophyre intrusions. acknowledgements the maps and field data presented here are the result of a 6-week expedition to tuttutooq in 2019. we acknowledge expedition support from terra nova, expedition foods, balfour beatty, veitchi group and blacks outdoor retail ltd. we also thank prof. brian upton for his insightful comments and discussions. place names are given in modern greenlandic orthography – older spellings present in the literature are tugtutôq (tuttutooq), sigsardlugtôq (sissarluttooq), kryds sø (krydssø), marrait (marraat) and ilímaussaq (ilimmaasaq). the older spelling has been retained for geological features named before the introduction of modern greenlandic orthography. additional information funding statement lk, rw, rc, lm and am received funding from the mining institute of scotland, institute of materials, minerals and mining, the edinburgh geological society, the augustine courtauld trust and the scott polar research institute. lk received funding from the society of economic geology hickok-radford fund. author contributions lk, rw, rc, lm and am: fieldwork, sample collection, drafting and editing of the manuscript. wm and af: research supervision and revision of the manuscript. competing interests the authors declare no competing interests. additional files a high-resolution geological map and a methods description can be found at https://doi.org/10.22008/fk2/nax2it references blaxland, a.b., van breemen, o., emeleus, c.h. & anderson, j.g. 1978: age and origin of the major syenite centers in the gardar province of south greenland: rb-sr studies. bulletin of the geological society of america 89(2), 231–244. https://doi.org/10.1130/0016-7606(1978)89<2 31:aaootm>2.0.co;2 cashman, k.v. & giordano, g. 2014: calderas and magma reservoirs. journal of volcanology and geothermal research 288, 28–45. https:// doi.org/10.1016/j.jvolgeores.2014.09.007 irvine, t. 1980: magmatic density currents and cumulus processes. american journal of science 280(a), 1–58. rosa, d., sandrin, a., nielsen, t.f.d. & vesturklett, h. 2020: petrography, geochemistry and magnetic susceptibility of the isortoq fe-ti-v deposit, isortoq giant dykes, south greenland. geus bulletin 44, 1–8. https://doi.org/10.34194/geusb.v44.4626 upton, b.g. 2013: tectono-magmatic evolution of the younger gardar southern rift, south greenland. geological survey of denmark and greenland bulletin 29, 1–24. https://doi.org/10.34194/geusb. v29.4692 upton, b.g.j. 1962: geology of tugtutoq and neighbouring islands, south greenland, part 1. meddelelser om gronland 169(8), 1–60. upton, b.g.j., craven, j.a. & kirstein, l.a. 2006: crystallisation of mela-aillikites of the narsaq region, gardar alkaline province, south greenland and relationships to other aillikitic-carbonatitic associations in the province. lithos 92(1–2), 300–319. https://doi.org/10.1016/j. lithos.2006.03.046 upton, b.g.j., emeleus, c.h., heaman, l.m., goodenough, k.m. & finch, a.a. 2003: magmatism of the mid-proterozoic gardar province, south greenland: chronology, petrogenesis and geological setting. lithos 68(1–2), 43–65. https://doi.org/10.1016/ s0024-4937(03)00030-6 upton, b.g.j. & thomas, j.e. 1980: the tugtutôq younger giant dyke complex, south greenland: fractional crystallization of transitional olivine basalt magma. journal of petrology 21(1), 167–198. https://doi. org/10.1093/petrology/21.1.167 waight, t., baker, j. & willigers, b. 2002: rb isotope dilution analyses by mc-icpms using zr to correct for mass fractionation: towards improved rb–sr geochronology? chemical geology 186(1–2), 99–116. https://doi.org/10.1016/s0009-2541(01)00420-x https://doi.org/10.34194/geusb.v47.6526 http://www.geusbulletin.org https://doi.org/10.22008/fk2/nax2it https://doi.org/10.1130/0016-7606(1978)89 https://doi.org/10.1016/j.jvolgeores.2014.09.007 https://doi.org/10.1016/j.jvolgeores.2014.09.007 https://doi.org/10.34194/geusb.v44.4626 https://doi.org/10.34194/geusb.v29.4692 https://doi.org/10.34194/geusb.v29.4692 https://doi.org/10.1016/j.lithos.2006.03.046 https://doi.org/10.1016/j.lithos.2006.03.046 https://doi.org/10.1016/s0024-4937(03)00030-6 https://doi.org/10.1016/s0024-4937(03)00030-6 https://doi.org/10.1093/petrology/21.1.167 https://doi.org/10.1093/petrology/21.1.167 https://doi.org/10.1016/s0009-2541(01)00420-x new insights from field observations of the younger giant dyke complex and mafic lamprophyres of the abstract introduction geological background results younger giant dyke complex ygdc facies description ygdc facies relationships marraat pod (l1) sissarluttooq pod (l2) krydssø pod (l3) assorutit pod (l4) ultramafic lamprophyre eastern lamprophyre and breccia western lamprophyre discussion origin of layering petrogenesis of ygdc magmas assorutit and krydssø magmas mafic enclaves homogeneous sheet timing of lamprophyre magmatism conclusion acknowledgements additional information references figures fig. 1 composite map of the studied area, with location identified by a red arrow in the inset map. fig. 2 selected photographs of key exposure within the ygdc. a: granite xenolith incorporated in the fig. 3 a: contact between a ‘homogenous finger’ and foliated troctolite at sissarluttooq. b: modal l fig. 4 focus maps of the western and eastern lamprophyre bodies (see fig. 1 for locations). table table 1 gardar intrusions on tuttutooq geological survey of denmark and greenland bulletin 7, 2004, p 9-12 9 although it was for many years believed that coals could not act as source rocks for commercial oil accumulations, it is today generally accepted that coals can indeed generate and expel commercial quantities of oil. while hydrocarbon generation from coals is less well understood than for marine and lacustrine source rocks, liquid hydrocarbon generation from coals and coaly source rocks is now known from many parts of the world, especially in the australasian region (macgregor 1994; todd et al. 1997). most of the known large oil accumulations derived from coaly source rocks have been generated from cenozoic coals, such as in the gippsland basin (australia), the taranaki basin (new zealand), and the kutei basin (indonesia). permian and jurassic coal-sourced oils are known from, respectively, the cooper basin (australia) and the danish north sea, but in general only minor quantities of oil appear to be related to coals of permian and jurassic age. in contrast, carboniferous coals are only associated with gas, as demonstrated for example by the large gas deposits in the southern north sea and the netherlands. overall, the oil generation capacity of coals seems to increase from the carboniferous to the cenozoic. this suggests a relationship to the evolution of more complex higher land plants through time, such that the highly diversified cenozoic plant communities in particular have the potential to produce oil-prone coals. in addition to this overall vegetational factor, the depositional conditions of the precursor mires influenced the generation potential. the various aspects of oil generation from coals have been the focus of research at the geological survey of denmark and greenland (geus) for several years, and recently a worldwide database consisting of more than 500 coals has been the subject of a detailed study that aims to describe the oil window and the generation potential of coals as a function of coal composition and age. depositional conditions hydrocarbons are derived from the aliphatic chains in the organic matrix. the dominant organic matter in conventional marine source rocks (type ii kerogen), and also lacustrine source rocks (type i kerogen), is formed from algae. this type of organic matter is geochemically quite uniform and contains an abundance of long-chain n-alkanes, a prerequisite for oil formation. coals are composed of transformed higher land plant material (mainly vitrinite or type iii kerogen) and are compositionally much more heterogeneous and complex. compared to algae-derived organic matter, the organic matter is richer in oxygen and contains fewer longchain n-alkanes. therefore coals inherently have lower oil generation potential. however, paralic coals that have been influenced by seawater during deposition may be enriched in hydrogen (higher h/c ratios; figs 1, 2; petersen & rosenberg 1998; sykes 2001), and incorporation of the hydrogen into aliphatic chains may increase the generation capacity. hydrogen enrichment is related to the activity of sulphate -reducing bacteria, and is commonly associated with increased contents of sulphur (fig. 1). in the søgne basin (north sea), for example, a clear facies-related change in generation potential is recorded for middle jurassic coal source rocks. the coals were formed on marine-influenced coastal geological survey of denmark and greenland bulletin 7, 9–12 (2005) © geus, 2005 oil generation from coal source rocks: the influence of depositional conditions and stratigraphic age henrik i. petersen c en o zo ic ju ra ss ic pe rm ia n c ar b o n ife ro u s > 1% sulphur (d.a.f.) 0.5–1% sulphur (d.a.f.) < 0.5% sulphur (d.a.f.) 0 0.1 0 0.4 0.2 0.8 0.6 1.4 type i type ii type iii 1.2 1 h /c ( at o m ic r at io ) o/c (atomic ratio) 0.2 0.3 fig. 1. the h/c ratio of coals from the major coal-forming geological periods displayed on the so-called ‘van krevelen diagram’. note that coals with a low sulphur content have a tendency to plot towards the lower part of the ‘type iii’-band and coals with a higher sulphur content towards the upper limit of the band. this feature is not related to coal age; d.a.f., dry ash-free. plains, and the thickest, most landward parts of the coals generated gas and condensate (harald field and trym discovery), whereas the coals that formed close to the palaeo-coastline generated oil (lulita field; e.g. petersen et al. 2000; petersen & brekke 2001). the influence of depositional conditions on the generation potential is independent of coal age (fig. 2; petersen in press). the depositional conditions of peat-forming mires may thus help to predict the source rock quality of the coals. generation potential the evolution of the generation potential (hydrogen index = hi) with increasing maturity is shown in fig. 2. the majority of the coal samples fall within a band that narrows with increasing maturity due the gradual homogenisation of the organic matter. the generation potential is exhausted around a vitrinite reflectance of 2.0% ro, but the coals may still possess a considerable potential (hi up to 190 mg hc/g toc) at the end of the conventional oil window at about 1.3% ro. a prominent feature of the hi-band is the initial increase in hi up to a maximum hi-value (e.g. huc et al. 1986; sykes 2001; petersen 2002, in press), defined by the himax line between 0.6% ro and 1.0% ro. the increase in hi is caused by the formation of an additional generation potential due to structural reorganisation of the coal matrix, possibly including incorporation of water-derived hydrogen (lewan 1997; schenk & horsfield 1998). this means that the ‘true’ generation potential of coals is equivalent to the himax, which is derived by translating the coals along their maturation pathway to the maximum value (fig. 2; sykes & snowdon 2002). coals with an himax < 150 mg hc/g toc are considered to be mainly gas-prone (fig. 2). compared to carboniferous, permian and jurassic coals, cenozoic coals attain the highest himax values (himax = 250–370 mg hc/g toc; petersen in press). however, the himax value may not necessarily be an expression of the ability to generate oil. the type of generated petroleum is dependent on the chain length of the n-alkanes in the coal structure. the hi is a measure of the hydrogen in the coal, but this may not be present as long-chain n-alkanes, but rather as shorter chains with a potential to form only gas or condensate. this may particularly be the case for carboniferous coals. coals with approximately similar hi values should theoretically contain the same proportion of hydrogen and thus have potentially the same petroleum generation capacity. information about the type of generated petroleum can be obtained by investigating the chain length of the n-alkanes in the solid coal structure by fourier transform infrared spectrometry (ftir) and by ruthenium tetroxide catalysed oxidation. the peak at 2850 cm–1 in the ftir spectrum is used to estimate the relative proportion of ch2 (dry, ash-free basis), which is taken as a measure of the proportion of oliphatic chains in the coal (fig. 3). in general, carboniferous coals contain a lower proportion of ch2 compared to cenozoic and jurassic coals with similar hi values. similarly, the proportion of aliphatic hydrogen is lower for carboniferous coals (fig. 3). cenozoic coals also have significant gas potential (estimated from the ch3 peak at 2955 cm–1), which is in good agreement with the coal-sourced south-east asian oil fields, that commonly also contain significant proportions of gas (e.g. the kutei basin, indonesia). ruthenium tetroxide catalysed oxidation of the coal structure essentially ‘chops’ off the aliphatic chains in the coal matrix, which after appropriate chemical treatment can be analysed by gas chromatography mass spectrometry (gc-ms). by adding an internal standard the obtained chromatograms can be directly compared. in general, carboniferous coals contain low proportions of n-alkanes with a carbon number > c19, whereas cenozoic coals in particular are much richer in long-chain aliphatics. the difference in the chemical structure between, for exam10 fig. 2. evolution in the hydrogen index (hi) with increasing maturity shown by 507 coal samples (petersen in press). coals with an hi < 150 mg hc/g toc are considered mainly gas-prone (shaded area). coals with increased sulphur contents have a tendency for higher hi values due to hydrogen-enrichment in the coal-structure. during initial thermal maturation the hi increases to a maximum value (himax), which is a better estimate of the true generation potential of a coal. the himax can be estimated by translating the hi value for a coal along its maturation pathway to the himax line (sykes & snowdon 2002); for example the coal with an hi of 185 (blue star) shows a correction of hi by 50 mg hc/g toc (himax = 230 mg hc/g toc); d.a.f., dry ash-free. 0 0.5 1.51 2 2.5 3 c en o zo ic ju ra ss ic pe rm ia n c ar b o n ife ro u s > 1% sulphur (d.a.f.) 0.5–1% sulphur (d.a.f.) < 0.5% sulphur (d.a.f.) 500 400 300 200 100 185 235 0 ga sp ro n e o ilp ro n e h yd ro ge n i n d ex ( m g h c /g t o c ) vitrinite reflectance (%ro) ~0.60%ro ~1.0%ro line of himax h h ple, carboniferous and cenozoic coals, can be related to the original coal-forming vegetation, and may explain why carboniferous coals are principally gasor condensate-prone, whereas cenozoic coals can be highly oil-prone. in contrast to the carboniferous woody, gymnospermous mire vegetations, the cenozoic coals were formed from an advanced and diverse vegetation, which may have produced a more aliphatic-rich vitrinitic organic matter upon deposition. the effective oil window (oil expulsion window) the complex, heterogeneous composition of coals results in a three-phase oil generation model (fig. 4; petersen in press): (1) onset of hydrocarbon generation, (2) hydrocarbon buildup in the coal to the expulsion threshold, and (3) oil expulsion in the so-called effective oil window. figure 4 shows the free hydrocarbons in the coals with increasing maturity. from 0.6–0.7% ro to 0.85–1.0% ro, the amount of oil increases in the coals up to a maximum value, after which it decreases. this decrease indicates the onset of efficient oil expulsion, and the maximum bi-value thus corresponds to the start of the effective oil window (oil expulsion window). the maturity at which oil expulsion starts and the range of the effective oil window is dependent on the initial generation potential of the coals (fig. 4). it should, however, be noted that the effective oil window for coals extends to higher maturities than the conventional oil window. coals generating < 12 mg hc/g 11 fig. 4. evolution in bitumen index (bi = s1 / toc) with increasing maturity. initially, hydrocarbons build up in the coal to a maximum value, which indicates the start of the effective oil window. coals with < 12 mg hc/g toc (shaded area) are considered to have a limited expulsion efficiency (gas-prone). in general, cenozoic coals generate the highest amounts of hydrocarbons, and they also reach the expulsion threshold at the lowest maturities as shown by the extension of the expulsion line down to 0.65% ro. (a) carb. (b) m. jurassic (a) carb. (b) m. jurassic3000 2950 28502900 2800 2750 cm-1 carboniferous coal, ruhr, germany m. jurassic coal, danish north sea ch2 ch3 hal/h (wt.%, d.a.f.) hal/mg coal (d.a.f.) 2850 cm-1: symmetric ch2 2955 cm-1: asymmetric ch3 h/c = 0.76 hi = 182 0.83%ro h/c = 0.79 hi = 205 0.82%ro a b c 60 40 20 0 b it u m en i n d ex ( m g h c /g t o c ) ~0.65%ro ~0.85%ro ~1.0%ro ~1.05%ro gas-prone 0 0.5 1.51 2 2.5 3 vitrinite reflectance (%ro) start of the effective oil window (oil expulsion window) unknown age cenozoic jurassic permian carboniferous fig. 3. a: ftir spectra (aliphatic stretching region) of carboniferous and middle jurassic coals with similar h/c ratios, hi values, and vitrinite reflectances. the middle jurassic coal shows a much higher relative response at 2850 cm–1, which is an indication of a higher proportion of aliphatic chains in the coal structure. b: calculated relative proportions of aliphatic chains (ch2) and ch3. c: calculated relative proportions of aliphatic hydrogen. the low values for the carboniferous coal suggest a primary gas generation potential; d.a.f., dry ash-free. 12 toc are considered to possess a limited oil expulsion efficiency. cenozoic coals generate the largest amounts of hydrocarbons and they tend to reach the expulsion threshold at the lowest maturities (from 0.65% ro; fig. 4). the broadest effective oil window is thus related to cenozoic coals, and may extend from approximately 0.65–2.0% ro. jurassic, permian and carboniferous coals reach the start of the effective oil window at higher maturities (0.85–0.9% ro). the jurassic coals show some ability to generate and expel hydrocarbons, whereas several of the carboniferous and permian coals seem to have a limited expulsion efficiency; these coals are principally gas-prone. as described above, the organic matter in carboniferous coals possesses an inherently low ability to generate oil. hence, the effective oil window for the carboniferous coals is in reality an effective gas/condensate window. conclusion coals can act as source rocks for oil accumulations. the generation potential is related to the depositional conditions of the coal-forming mires, with marine influence having a positive effect by increasing the hydrogen content in the vitrinitic organic matter. in addition, an overall vegetational control seems to be exerted on the source rock potential. carboniferous coals contain lower proportions of long-chain n-alkanes and aliphatic hydrogen in the coal matrix than younger coals. cenozoic coals generally contain high proportions of long-chain n-alkanes and possess a high oil generation potential in addition to a high gas potential. during maturation, coals form an additional generation potential, and hydrocarbon generation can be described as a three-phase process, including a hydrocarbon build-up phase in the coals before onset of efficient expulsion. the effective oil window starts at higher maturities than the conventional oil window, and in addition extends to higher maturities. cenozoic coals display the broadest effective oil window, whereas the effective oil window for carboniferous coals is, in reality, an effective gas/condensate window. this enhanced understanding of oil generation from coal source rocks is being directly employed by geus in petroleum geological projects that are being carried out, for example, in vietnam (cf. nielsen & abatzis 2004). acknowledgement the carlsberg research foundation (ans-1293 & ans-1293/20) is thanked for financial support. references huc, a.y., durand, b., roucachet, j., vandenbrouke, m. & pittion, j.l. 1986: comparison of three series of organic matter of continental origin. organic geochemistry 19, 191–204. lewan, m.d. 1997: experiments on the role of water in petroleum formation. geochimica et cosmochimica acta 61, 3691–3723. macgregor, d.s. 1994: coal-bearing strata as source rocks – a global review. in: scott, a.c. & fleet, a.j. (eds): coal and coal-bearing strata as oil-prone source rocks? geological society special publication (london) 77, 107–116. nielsen, l.h & abatzis, i. 2004: petroleum potential of sedimentary basins in vietnam: long-term geoscientific co-operation with the vietnam petroleum institute. geological survey of denmark and greenland bulletin 4, 97–100. petersen, h.i. 2002: a re-consideration of the ‘oil window’ for humic coal and kerogen type iii source rocks. journal of petroleum geology 25, 407–432. petersen, h.i. in press: the petroleum generation potential and effective oil window of humic coals related to coal composition and age. international journal of coal geology. petersen, h.i. & brekke, t. 2001: source rock analysis and petroleum geochemistry of the trym discovery, norwegian north sea: a middle jurassic coal-sourced petroleum system. marine and petroleum geology 18, 889–908. petersen, h.i. & rosenberg, p. 1998: reflectance retardation (suppression) and source rock properties related to hydrogen-enriched vitrinite in middle jurassic coals, danish north sea. journal of petroleum geology 21, 247–263. petersen, h.i., andsbjerg, j., bojesen-koefoed, j.a. & nytoft, h.p. 2000: coal-generated oil: source rock evaluation and petroleum geochemistry of the lulita oilfield, danish north sea. journal of petroleum geology 23, 55–90. schenk, h.j. & horsfield, b. 1998: using natural maturation series to evaluate the utility of parallel reaction kinetics models: an investigation of toarcian shales and carboniferous coals, germany. organic geochemistry 29, 137–154. sykes, r. 2001: depositional and rank controls on the petroleum potential of coaly source rocks. in: hill, k.c. & bernecker, t. (eds): eastern australasian basins symposium, a refocused energy perspective for the future. petroleum exploration society of australia special publication, 591–601. sykes, r. & snowdon, l.r. 2002: guidelines for assessing the petroleum potential of coaly source rocks using rock-eval pyrolysis. organic geochemistry 33, 1441–1455. todd, s.p., dunn, m.e. & barwise, a.j.g. 1997: characterizing petroleum charge systems in the tertiary of se asia. in: fraser, a.j., matthews, s.j. & murphy, r.w. (eds): petroleum geology of southeast asia. geological society special publication (london) 126, 25–47. author’s address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: hip@geus.dk geological survey of denmark and greenland bulletin 12, 57-63 57 pyritic part of the knudshoved member may correlate with theupper, tuff-poorunitb2 of the balder formation. stronsay group knox & holloway (1992) replaced the hordaland group of deegan & scull (1977) with two new groups: the stronsay group succeeded by the westray group (fig. 3). the two groups together comprise the light grey, green and brown coloured, soft, fissile, marine shales with thin limestone streaks that overlie the rogaland group and underlie the nordland group. these groups each contain two formations, one representing sandy shelf lithofacies and the other representing basinal mudstone lithofacies. in the central north sea, and in the danish sector, the stronsay group is represented by its mudstone facies, the horda formation (knox & holloway 1992). sandstone units of varying thickness occur at many levels in the stronsay and westray groups along the basin margin in the norwegian and british sectors, and many of these have been defined as formations or members (deegan & scull 1977; hardt et al. 1989; knox & holloway 1992). a sandstone unit also occurs in the horda formation on the ringkøbing–fyn high in the danish sector and is described here as a new member (hefring member). horda formation history. knox & holloway (1992) established the horda formation for the greenish grey basinal mudstone facies of their stronsay group that overlies the grey tuffaceous mudstones of the balder formation and underlies the greenish grey to brown mudstones of the lark formation (knox & holloway 1992). type well. british sector well 22/1-1a, 2379.5–1992 m mdkb. danish reference wells. mona-1, 2930.8–2363.5 m mdkb (fig. 46); siri-1, 2037.9–1916.5 m mdkb (fig. 47). distribution and thickness. the horda formation extends over the central and northern north sea and is present in 1916.5 2037.9 horda fm chalk gp lark fm balder fm sele fm lista fm våle fm 1900 2000 2100 2200 m siri-1 gr sonic rind mb idun mb tyr mb vile mb ve mb bue mb bue mb horda formation siri-1 mona-1 floki-1 100 200 300 400 500 600 700 800 900 thickness (m) 25 km fig. 47. siri-1, danish reference well for the horda formation. black bars show cored sections. fig. 48. isochore map of the horda formation in the study area. the positions of the two danish reference wells for the horda formation, mona-1 and siri-1, are indicated on the map. the position of floki-1, the type well for the hefring member, is also indicated. 58 gulnare-1 adda-2 alma-1 nnw sse 10 km horda fmhorda fmhorda fm l4l4l4 l2l2l2 0 500 1000 1500 2000 2500 3000 3500 twt (msec) tl tf uou tl2 th tb tc l3l3 lark fmlark fmlark fm 1000 1500 2000 2500 3000 3500 gert-1 sw ne mona-1 elna-1 sandra-1 horda fmhorda fmhorda fm h2h2 h3h3th2th2 h2 h3h3th2 10 km lark fmlark fmlark fm tl uou th tb tc twt (msec) th1th1th1 h1h1h1 fig. 49. nnw–sse-trending seismic section (rtd81-re94-17a) in the central graben showing southand eastward thinning of the horda formation. the location of the seismic section is shown in fig. 1. the l2, l3 and l4 subunits of the lark formation are indicated, as well as the mudstone-equivalent of the freja member. tl, top lark; tf, top freja; uou, upper oligocene unconformity; tl2, top l2; th, top horda; tb, top balder; tc, top chalk. fig. 50. sw–ne-trending seismic section (rtd81-rtd94-19a) showing the tripartite subdivision (h1–3) of the horda formation in the eastern part of the danish central graben (gert-1 and mona-1) and pronounced thinning of the horda formation east of the central graben (elna-1 and sandra-1). the location of the seismic section is shown in fig. 1. th1, top h1 marker; th2, top h2 marker; other abbreviations as in fig. 49. 59 all wells in the danish sector of the north sea. however, the lower part of the horda formation (fig. 4; equivalent to sequence 2 of michelsen et al. 1998) is lacking in the eastern wells r-1 and s-1 and in the eastern part of the ringkøbing–fyn high. the upper part of the horda formation (fig. 4; equivalent to sequence 3 of michelsen et al. 1998) is thin or absent in the same area (michelsen et al. 1998). the horda formation reaches a thickness of 906 m in the central graben well tordenskjold-1, but thins towards the east and south-east to less than 100 m, with minimum recorded thicknesses of 9 m in the ida-1 well and 4 m in the s-1 well. an isochore map of the horda formation is shown in fig. 48. the overall thinning of the horda formation towards the south-east, east and north-east is also shown on the seismic sections in figs 49 and 50 and on the log panel in fig. 51. lithology. the horda formation is characterised by greenish grey to greyish green fissile mudstone. subordinate limestone benches and thin layers of black mudstones occur at some levels in the formation. in many wells, particularly in the central graben, the lowermost 20–50 m of the horda formation consists of red-brown mudstones (fig. 52). this lithology is apparently lacking in the eastern wells of the danish sector. log characteristics. the horda formation is characterised by an overall stable gamma-ray and sonic log motif with a lower gamma-ray response than that displayed by the underlying balder formation and the overlying lark formation. in a few wells, the base of the horda formation shows relatively high gamma-ray values, which decrease to lower and more stable values over a short interval. the sonic readings decrease slightly upwards from the base to the top of the horda formation. boundaries. the base of the horda formation is placed at the change from the laminated, predominantly grey mudstones with interbedded sandy tuffs of the balder formation to the predominantly non-laminated, fissile, greenish grey or red-brown massive mudstones that form the basal part of the horda formation. the balder–horda boundary may be conformable or marked by a hiatus. the boundary is often difficult to pick on petrophysical logs. in basinal settings, knox & holloway (1992) advocated placing the lower boundary of the horda formation at the base of a marked gamma-ray peak believed to represent a glaucony-rich condensed layer in the basal part of the horda formation. however, in many sections in the danish sector there are two or more gamma-ray peaks in the balder–horda boundary interval. as the glauconyrich layer has not been identified with certainty in the few cores taken across the boundary in the danish sector, it is not possible to identify the key gamma-ray peak unambiguously. therefore, it is suggested that the lower boundary of the horda formation is placed on the basis of the sonic log where a gradual decrease in values in the upper part of the balder formation is succeeded by relatively stable, but somewhat lower readings in the horda formation (figs 46, 47). the upper boundary is at the base of the lark formation. subdivision. knox & holloway (1992) suggested a threefold subdivision of the horda formation (h1–3), based on lithology and biostratigraphy. a threefold subdivision can also be seen on seismic sections in the danish sector of the north sea (fig. 50). in some central graben wells, the subdivision may also be recognised on shifts in log patterns on both gamma-ray and sonic logs (fig. 51). in these wells, subtle peaks separate the three units on the gamma-ray log and coincide with the top h1 and top h2 seismic markers. based on analysis of cuttings samples, the subdivision apparently lacks lithological expression in the danish sector. knox & holloway (1992) noticed that the top of unit h1 is close to the ho of the dinoflagellate eatonicysta ursulae, and that the top of unit h2 is close to the ho of the foraminifer spiroplectammina spectabilis. this observation is supported by biostratigraphical data from the present study. a sandstone body within the horda formation has been encountered in the well floki-1 on the ringkøbing–fyn high) in the danish sector of the north sea. this sandstone is defined herein as the new hefring member (see below). macroand ichnofossils. the horda formation is moderately to intensely bioturbated. ichnofossils comprise chondrites ispp., phycosiphon ispp. and planolites ispp. microfossils and palynomorphs. in wells where the horda formation rests conformably on the balder formation, the dinoflagellate events ho deflandrea oebisfeldensis and ho dracodinium condylos occur in the lowermost part of the horda formation. characteristic microfossil datums from the lowermost part of the horda formation are the ho of the planktonic foraminifer subbotina ex gr. linaperta, which occurs abundantly, followed upwards by the almost coeval hos of the benthic foraminifers uvigerina batjesi, turrilina brevispira and gaudryina hiltermanni. a hiatus between the horda and balder formations is indicated in wells in the northern and eastern part of the danish sector by the absence of d. oebisfeldensis and d. condylos from the lower part of the horda formation. central 60 1700 1800 1900 1400 gr sonic gr sonic density gr sonic density gr sonic density 1500 1600 2000 2100 2200 2300 2400 2500 2600 2700 2800 2900 3000 3100 m 1000 1100 1200 700 800 900 1300 1400 1500 1600 1700 1800 1900 2100 2200 m 1700 1800 1900 1400 1500 1600 2000 2100 2200 2300 2400 2500 2600 2700 2800 2900 m 1800 2000 1500 2100 2200 2300 2400 2500 2600 2700 2800 3000 m kim-1 kim-1 mona-1 west lulu-3 siri-1 siri-1 1600 1700 1900 2900 2000 horda h1 h3 h2 lark l4 tl uou tl2 tl1 th th2 th1 tb l3 l2 l1 balder sele lista våle chalk grp sw ne 61 above the top of the lower, oxic part of the formation. the occasional influxes of radiolaria recorded throughout the upper part of the formation suggest that deeper marine conditions prevailed periodically. the palynofacies of the horda formation is characterised by a rich and dominant dinoflagellate assemblage with dispersed terrestrial matter (phytoclasts, spores and polfacing page: fig. 51. sw–ne-trending log panel showing eastward thinning of the horda formation. the figure also shows the variation in architecture and distribution of the horda units h1–h3 and lark units l1– 4 between the central graben (kim-1, mona-1 and west lulu-3) and the ringkøbing–fyn high (siri-1). seismic ties: tl1, top l1; th2, top h2; th1, top h1; other abbreviations as in fig. 49. fig. 52. core photographs showing red to reddish grey mudstones of the lowermost part of the horda formation in the sofie-1 well. depths are core depths. cm 1887 m 1888 m sofie-1 1901 m 1902 m 0 10 20 30 40 50 60 70 80 90 100 graben wells contain the downhole succession of the dinoflagellate cyst events ho areosphaeridium diktyoplokum, ho areosphaeridium michoudii, ho heteraulacacysta porosa and ho cerebrocysta bartonensis from the top of the horda formation, indicating an age as young as earliest rupelian (fig. 5b). in wells to the north and east, the top of the horda formation seems to be slightly older since a. michoudii is recorded from the top of the formation, indicating a mid-priabonian age (fig. 5b). significant dinoflagellate events from the middle to upper part of the horda formation are the succession of the hos of eatonicysta ursulae, diphyes ficosoides and phthanoperidinium clithridium in the middle part of the formation, and the hos of diphyes colligerum and c. bartonensis in the upper part of the formation. in central graben wells, significant microfossil events in the lower part of the horda formation include the ho of abundant radiolaria of the genus cenosphaera and the ho of the planktonic foraminifer cyclammina amplectens. key events in the middle and upper parts of the formation are the hos of pseudohastigerina spp. (planktonic foraminifers), lenticulina gutticostata, spiroplectammina amplectens and planulina costata (benthic foraminifers). the top of the horda formation contains the hos of cibicidoides truncanus and vaginulinopsis decorata. depositional environment. the lower part of the horda formation contains a microfauna that is significantly different from that of the underlying balder formation. the basal 5–40 m of the horda formation are characterised by a diverse fauna of both benthic and planktonic calcareous foraminifers together with agglutinated foraminifers. this indicates that the depositional setting was open marine, bathyal and with oxic bottom conditions. the upper part of the horda formation is characterised by an abundant and diverse agglutinated foraminifer fauna. calcareous foraminifers are very sparse or absent in this interval. the assemblage of rhabdammina discreta, haplophragmoides spp., recurvoides spp. and usbekistania charoides indicate that the upper part of the horda formation was deposited at upper bathyal depths with dysoxic bottom conditions. radiolaria occur commonly in several narrow intervals, the lowest of which is slightly 62 horda fm lark fm hefring mb våle fm balder fm sele fm lista fm chalk gp 1500 1600 1700 1800 m floki-1 1731.3 1793.4 gr sonic neutron/density fig. 53. floki-1, type well for the hefring member. len) as a minor component, indicating an open marine environment with only limited influx from surrounding terrestrial areas. age. in the central graben, where the horda formation is most complete, the formation spans from the middle ypresian (early eocene) at its base to earliest rupelian (earliest oligocene) at its top. in wells to the east and north, the top is as old as middle priabonian (see also biostratigraphic section above). this indicates that the top of the horda formation is diachronous, younging in a southwesterly direction. this is possibly due to increased erosion or longer intervals of non-deposition towards the north-east in the basin, or both. correlation. the horda formation can be correlated with the onshore danish succession of the røsnæs clay formation, the lillebælt clay formation and the søvind marl formation (heilmann-clausen et al. 1985), and the viborg formation (christensen & ulleberg 1973). the red-brown mudstones near the base of the horda formation in the central north sea can be correlated lithologically with the røsnæs clay formation and the lower part of the lillebælt clay formation. 63 the lower part of the overlying main body of greenish and greyish mudstones in the offshore succession can be correlated with the coeval and lithologically similar upper part of the lillebælt clay formation. the upper part of the horda formation can be correlated with the søvind marl formation, which consists of grey marls. the highest part of the horda formation, only observed in central graben wells, may be correlated with the viborg formation on biostratigraphic evidence. hefring member new member history. the hefring member consists of sandstone deposits within the horda formation. these sandstones have not previously been recognised as a separate unit in the danish sector. derivation of name. after the goddess hefring. type well. danish sector well floki-1, 1793.4–1731.3 m mdrt (fig. 53). distribution and thickness. the hefring member is only known from the floki-1 well located in the northern part of the danish sector. as the unit currently cannot be identified on seismic sections, its further distribution is unknown. in the floki-1 well, the member is 62 m thick. lithology. the hefring member consists of greenish grey, fine-grained, immature sandstones with glaucony grains. logcharacteristics.thehefring member ischaracterisedby aconspicuousblockysignatureonthegamma-ray, sonicand density logs (fig. 53). gamma-ray responses are lower than those of the enveloping horda formation mudstones. the hefring member can also be recognised from a combination of the density and neutron logs as the presence of pure sandstones results ina ‘cross-over’of the two logcurves (fig. 53). boundaries. the boundaries with the mudstones of the horda formation are sharp and characterised by prominent shifts on the gamma-ray and sonic logs (fig. 53). depositional environment. no cores have been taken in the hefring member, but the sandstones were probably deposited from concentrated gravity flows, based on log similarity with the other fine-grained sandstone bodies in the nearby siri canyon. age. lutetian (middle eocene) based on the age of the associated horda formation mudstones. correlation. based on biostratigraphic data, the hefring member may be contemporaneous in part with the lillebælt clay formation onshore denmark, with the lower part of the grid sandstone member (knox & holloway 1992) in the viking graben and with the upper part of the tay sandstone member (knox & holloway 1992) in the northern part of the central graben. westray group the westray group is the upper of the two groups established by knox & holloway (1992) to replace the hordaland group of deegan & scull (1977; fig. 3). in the central north sea and in the danish sector of the north sea, the westray group is represented by the lark formation. lark formation history. the lark formation was established by knox & holloway (1992) for the brownish grey mudstone-dominated lithofacies of the westray group that overlies the more variable association of red and green-grey mudstones, silty mudstones and sandstones of the horda formation and underlies the grey, sandy and shelly mudstones, siltstones and sandstones of the nordland group of deegan & scull (1977; fig. 3). the lark formation is also recognised in the danish sector although its lithology is more variable than that given in the original description. type well. british sector well 21/10-4, 1867–1217 m mdkb. danish reference wells. mona-1, 2363.5–1598.3 m mdkb (fig. 46); siri-1, 1916.5–819.3 m mdkb (fig. 54). distribution and thickness. the lark formation extends over the central and northern north sea and is probably present in the entire danish sector of the north sea. its depocentre is in the central and northern part of the danish sector, along the eastern boundary of the danish central graben, where it reaches a thickness of 1194 m in the siri-3 well. the lark formation thins west to a thickness of 389 m in the tordenskjold-1 well in the central graben, and east to a thickness of 240 m in the s-1 well on the ringkøbing–fyn high (fig. 55). descriptive text to the geological map of denmark, map description pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 1 of 53 descriptive text to the geological map of denmark, 1:50 000, møn 1511 i, 1511 iv and 1512 ii stig a. schack pedersen*1 , peter gravesen1 1geological survey of denmark and greenland (geus), copenhagen, denmark. abstract the geological map sheet møn covers the island of møn, the smaller adjacent islands langø, lindholm and nyord as well as adjacent parts of sjælland and lolland. it comprises the geodetic map sheets 1511 i and 1511 ii and areas on bordering sheets. møn is surrounded by the baltic sea with the bay of hjelm bugt to the south, the straits of grønsund and ulvsund to the west, and the bays of stege bugt and fakse bugt to the north. møn is divided into three glaciomorphological areas, namely a high, hilly landscape of høje møn to the east, a hummocky to parallel ridge landscape to the west and areas of marine deposits around nyord and ulvshale. the composite ridge landscape of høje møn constitutes a glaciotectonic complex comprising four individual glaciodynamic sequences, with the hill aborrebjerg as the highest point (143 m a.s.l.). the parallel ridge hills consist of thrust-fault-displaced chalk sheets with superimposed glacial deposits. the thrust sheets are up to 80 m thick, of which 60 m constitute maastrichtian chalk. the vertical displacement of the thrust sheets is about 150 m measured from the primary, undeformed pre-quaternary surface located 25–30 m below sea level. the pre-quaternary surface consists of chalk of late maastrichtian age, which forms a carbonate platform in the subsurface of møn about 27 m below sea level. chalk displaced by glacial tectonics is not restricted to høje møn but also appears in smaller thrust sheets and rafts in the small-ridged landscape around stege nor. in the chalk sheets along møns klint, most of the late maastrichtian succession is exposed. cliff sections with chalk are also exposed at hvideklint along the south coast of the island. however, here the glaciotectonic shear deformation has commonly altered the lithology into a chalk glacitectonite. the oldest quaternary units deposited on the pre-quaternary unconformity are saalian till as well as sand and clay from the eemian interglacial. these units are overlain by early weichselian sand. the next quaternary succession, the ristinge klint till formation, was deposited during the ristinge ice advance in the early middle weichselian about 55 000–50 000 years ago. then followed the kraneled formation (new formation) consisting of fluvial and lacustrine deposits. the following klintholm till formation (adjusted formation) was deposited during the klintholm ice advance 35 000–32 000 years ago. the klintholm till formation is overlain by a more than 10 m thick unit of greyish glaciolacustrine clay with dropstones. glaciofluvial sand with thin-layered intercalations of laminated mud and diamictites of the kobbelgård formation (new formation) are related to this unit and interpreted as deposited in a huge, partly ice dammed lake covering a large part of the present baltic sea and the southern part of kattegat 32 000 to 28 000 years ago. the kobbelgård formation is overlain by sand and gravel of the stubberup have formation (new formation) and tills of the mid danish till formation deposited by the ne ice advance from central sweden about 23 000–20 000 years ago. relatively shortly after the ne ice had melted away, the young baltic ice advanced from the eastern part of the baltic area. north-directed compressive deformation during this advance created the glaciotectonic complex of møns klint including the new unit møns klint glaciodynamic sequence. in the southern part of the complex, a steeply inclined imbricated fan was formed; towards the foreland to the north, the thrust faults became gently dipping and the tip-zone of thrusting is located under the landslides at liselund. the composite ridges form a characteristic hilly landscape with elongate crests trending e–w. the curved coastline along hjelm bugt was formed by a glacial lobe, north of which a push moraine was built up. a number of spillways striking radially northward from the lobe were formed by meltwater discharged from its glacier ports, including the borre, maglemose and røddinge depressions. deposition of sand and gravel of the ny borre formation (new formation) took place at this time. during the advance of the young baltic ice over southern denmark to the eastern jutland stationary line, a relatively thin lodgement till of the lolland till formation was deposited, which is rich in chalk due to its truncation of the upthrusted chalk sheets. *correspondence: sasp@geus.dk received: 04 feb 2019 accepted: 18 apr 2021 published (in danish): 14 dec 2021 published (in english): 16 dec 2022 keywords: geological map, quaternary stratigraphy, cretaceous chalk, glacial tectonics, landslides abbreviations: bp: years before present osl: optically stimulated luminescence tl: thermoluminescence geus: geological survey of denmark and greenland lgm: late glacial maximum geus bulletin is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: adam a garde (geus, denmark) translated by: adam a garde, 2022. this work was originally published in danish (https://doi.org/10.34194/geusb.v48.8293) reviewed by: helena alexanderson (lund university, sweden) & nikolaj krogh larsen (university of copenhagen, denmark) funding: see page 49 competing interests: none declared additional files: see page 49 pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 2 of 53 www.geusbul let in.org towards the end of the weichselian glaciation c. 17 000 years ago, the young baltic ice melted back, leaving a residual ice cap in skåne from where a recessive ice advance towards south-west reached møns klint, resulting in superimposed glaciotectonic deformation. during the late weichselian, freshwater lakes in the hjelm, tøvelde and høje møn areas were filled by clay and gyttja, with deposition that continued into the holocene. during the holocene, the former spillways were turned into fjords during the atlantic transgression. marine deposits mirroring the littorina sea are thus found in maglemose and borre sømose. after the atlantic transgression had established a sea level more or less corresponding to that of today, accretion of marine forelands and formation of a spit system started. in particular, this is the case for the areas of ulvshale and nyord. at the same time, vegetation migrated out into the numerous fjords, and peat began to accumulate. the last phase of sedimentation is confined to the formation of beach ridges in the coastal areas, typically covered by aeolian dunes, as can be seen on the coast at klintholm havn and råbylille as well as along the north-eastern coast of ulvshale. 10°e 56°n kattegat skagerrak femer bælt baltic sea storebæ lt lolland la ng el an d falster møn ristinge klint fyn jyl land sjæl land sweden denmark 50 km ø resund fig. 1 index map of denmark with the position of the map sheet møn. introduction topography the geological map sheet møn at scale 1:50 000 comprises the northern part of the topographic map sheet 1511 i, the southern half of 1512 ii and a smaller part of 1511 iv (fig. 1). it covers the island of møn and the surrounding sea with the islands of nyord, lindholm and langø, as well as smaller parts of south-eastern sjælland and north-eastern falster. the map area is located within 11°15’–11°42’ e and 54°47’–55°04’ n. all place names mentioned in the text are found in the map figures here and the accompanying map sheet (available online as a supplementary file and as a supplement to the printed version.) the part on south-eastern sjælland covers a small triangular area extending from kalvehave in the north to sandvig kohave with the forest of viemose skov to the south, and westwards from kalvehave to langebæk. in the west, the area reaches an altitude of 46 m a.s.l. and consists of a morain plateau sloping gently towards the coasts, where holocene marine deposits of clay and sand occur in several places, mainly at kindvig hoved. a small part of the map sheet area on north-eastern falster next to grønsund north of korselitse østerskov and east of stubbekøbing consists of till and holocene clay deposits in noret. towards grønsund færgebro and http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 3 of 53 www.geusbul let in.org along the coast towards grønsund, the angular foreland also comprises marine deposits, mainly sand, gravel and beach cobbles of flint (klint et al. 2017). based on the topography, møn comprises three characteristic geomorphological terrains: (1) the high, hilly landscape of høje møn, (2) the glaciomorphological landscape around stege nor, and (3) the flat marine foreland around nyord and ulvshale. the highest part of møn is in store klinteskov, the woodland that characterises høje møn. aborrebjerg at 143 m a.s.l. forms a marked ridge above the aborresø lake. several other hills, including kongsbjerg, lollikebakke, lindebakker and store ørnebjerg, reach elevations above 130 m. lakes with the highest altitude of water table in denmark, almost 90 m a.s.l., are located here, namely aborresø, st. geddesø and hunesø. the most dramatic topography is found at dronningestolen with an almost vertical fall of 130 m to sea level. a n–s-trending lowland around borre sømose and busemarke mose with elevations down to only 1 m above sea level separates høje møn from the glaciomorphological landscape around stege nor with general elevations around 25 m and local hilly areas up to c. 40 m above sea level. møn’s surface is dominated by glacial weichselian sediments with a high content of chalk, apart from large areas with holocene deposits at kostervig, nyord, ulvshale, maglemose, borre sømose, busemarke mose, råbylille sø and råby sø. particularly høje møn with møns klint and hvideklint at hjelm bugt are examples of glaciotectonic cliffs with intercalations of large chalk sheets within till and meltwater deposits, with characteristic surface expressions particularly in the former area. all of the flat marine foreland at ulvshale lies between 1–3 m a.s.l. the highest elevations are elongate aeolian ridges deposited on beach ridges. the dammed area at klostervig and former fjord areas such as borre sømose and maglemose also belong to this landscape group. the southernmost point of møn at hårbølle havn constitutes a marine angular foreland. large parts of møn’s coasts, especially in the south and east, but also in the north-east, consist of modified cliffs separated by marine forland. the coast to the west and south from ulvshale is lobate, with the largest embayments at stege nor and fanefjord. hjelm bugt, fakse bugt and stege bugt, which partly encircle møn, are shallow marine areas with few incisions of deeper sounds such as grønsund between møn and falster (up to 23 m deep) and ulvsund between sjælland and møn (up to 18 m deep). there are only a few small islands in these waters; nyord and langø are the largest. methods the systematic geological mapping of the danish land surface is traditionally also seen as a mapping of the quaternary geology, because more than 90% of the land surface consists of quaternary deposits. the mapping is carried out by the geological survey of denmark and greenland (geus) by walking through the landscape with a hand auger. at intervals of c. 100 m, a soil sample is collected at 1 m depth, attached to a small groove at the tip of the spear. the sample thus represents the subsurface material below the soil horizon and any cultural layers. the nature of the collected sample is determined on location and noted on a field map. in this way, boundaries between different deposits can be determined in the field, and a sufficient number of observations within each mapped polygon is ensured. the sample spacing may commonly be closer than 100 m, especially where geological boundaries need to be established (gravesen et al. 2006; jakobsen et al. 2015). information from road cuttings, gravel pits, other digging activity and coast profiles is also used. in the field, about 40 different types of lithological units are distinguished and marked on a field map at scale 1:25 000 with their appropriate symbols. these preliminary field maps are digitised and stored in geus’ database and are accessible at www.geus.dk. they constitute the basis for the published geological map. the principles of establishing the different types of quaternary deposits are described in pedersen et al. (2015). in the following descriptions, each unit is assigned to a marine isotope stage (mis), following cohen & gibbard (2012). all age data presented in this map sheet description follow the systematics of houmark-nielsen et al. (2017). the data collected during the systematic geological surface mapping is stored in geus’ archives. seventy per cent of the land area shown on the map sheet is till, which amounts to 218 km2 on møn itself besides smaller areas on sjælland and eastern falster. the sandy–clastic glacial deposits are dominated by meltwater sand and gravel, which cover 11% of the land surface. the remaining part mainly comprise holocene marine and freshwater deposits (10 and 7%, respectively) besides chalk (1%) mainly on høje møn, as well as lakes and unmapped areas of towns, roads (1%). since around 1985, the cliffs on southern, eastern and northern møn have been subject to several studies including several unpublished msc dissertations. the results of raw material and groundwater investigations have been published in various reports. results from all these studies have been included in the present map sheet description. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 4 of 53 www.geusbul let in.org history of geological mapping on and around møn the first geological observations on møn were pub lished by pontoppidan (1768), who demonstrated the geological structure of the southern part of møns klint from vidskud via græderen and sommerspiret to dronningestolen in a wood carving (see section on møns klint). puggaard (1851, 1852) performed a meticulous measurement of the structures all along møns klint and used this to describe the geological structure of the entire island. puggaard fully supported a glaciotectonic mechanism for the formation of møns klint, but it was not generally accepted prior to johnstrup’s (1874) publication, which stated that the chalk sheets were formed by pressure from an e–w-directed baltic ice stream through the baltic sea. furthermore, johnstrup pointed out the correlation between rügen and møn, and concluded that at both sites the chalk sheets with their cover of till were pushed up and imbricated. k.a. grønwall suggested already in 1906 that a geological map of the entirety of møn should be produced. hintze (1937) published a map of høje møn, which retained a block-tectonic interpretation of the deformation structures in møns klint. all subsequent studies support a glaciotectonic interpretation (konradi 1973; berthelsen et al. 1977; pedersen 1988, 2000, 2011; houmark-nielsen 1994). hintze (1937) interpreted all depressions on his map, which had no outlet, as some sort of karst phenomena, although almost all were caused by melting of dead-ice bodies left behind when the young baltic ice retreated. milthers (1948) compiled a quaternary map of denmark at scale 1:320 000. however, on møn, he only differentiated between till landscapes, dammed marine deposits and beach ridge complexes. v. haarsted presented a geomorphological map and a map of quaternary deposits of møn at scale 1:20 000 at a meeting of the danish geological society, january 1955. an overview map in the minutes from the presentation contains a summary of the geological features, where elongate ridges are interpreted as ice stationary lines. valleys and eskers are also marked (haarsted 1956). a compilation of quaternary maps of denmark at scale 1:200  000 (pedersen 1989) included information from drilling and aerial photographs. this allowed a designation of høje møn as consisting of chalk, and a few chalk sheets were shown in the easternmost part of the island. the mapping for the present map sheet of møn was performed by several different geologists and at various times. v. haarsted, who was headmaster at the secondary school of holte gymnasium, carried out his mapping until 1959. v. münther and a.v. nielsen mapped smaller areas between 1966 and 1972. the present authors completed the mapping in 2004–2006, following the publication of a preliminary map of eastern møn (pedersen & gravesen 2005). pre-quaternary deposits late cretaceous (99.6–66.0 ma) chalk underlies møn’s pleistocene deposits. møn is located on the southern part of the ringkøbing–fyn high (sorgenfrei & buch 1964; sorgenfrei 1966: håkansson & pedersen 1992; surlyk et al. 2013). in this region, chalk and limestone deposits about 500 m thick were laid down, whereas the equivalent deposits in the danish basin to the north are up to 2000 m thick (surlyk et al. 2013). most of these deposits are chalk. relatively thin layers of chalk (<500 m thick) also occur in the german basin south of the ringkøbing–fyn high (surlyk et al. 2013). -20 -20 -20 -20 -10 -10 -10 -10 -30 30 -30 -30 10 10 50 80 0 0 0 0 0 n 5 m fig. 2 the top of chalk below the quaternary deposits, including chalk sheets dislocated by glacial tectonics. contour lines with 10 m intervals shown above and below sea level. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 5 of 53 www.geusbul let in.org there are no deep boreholes on møn. the nearest one is the 2.5 km deep ørslev-1 borehole on falster, drilled by the oil company gulf in 1967–1968 (nielsen & japsen 1991; geus 1997). the base of the chalk is located at 450 m depth. the deepest part of the borehole comprises carboniferous deposits, and seismic data suggest that crystalline basement occurs less than 3 km below surface. layers of rock salt, which forms pillows and small diapirs on lolland, occur at a depth of 1.5 km. late cretaceous chalk the chalk is exposed on møn intercalated with quaternary deposits in many coast profiles affected by glacial tectonics, but only occurs in situ in boreholes. borehole data show that the pre-quaternary surface is located 25–40 m below sea level in large parts of møn, forming a relatively planar surface that only locally reaches greater depths (see profiles on the map sheet and in houmark-nielsen 2003). its position below the imbricated chalk sheets at høje møn is poorly constrained but possibly deeper. north of hvideklint and in the area between elmelunde and borre, the chalk appears to reach sea level, but this is probably due to glaciotectonic uplift of chalk sheets (fig. 2). the chalk is a white, greyish white or yellowish, soft, sludgy calcareous rock (fig. 3), which almost exclusively consists of coccoliths, micrometre-sized plates a b fig. 3 pale grey chalk with black flint layers. a: folded flint layers in the chalk cliff of st. stejlbjerg, southern møns klint. the chalk cliff is here almost 100 m high. b: close-up of black flint layers in chalk. the large, circular flint concretion in the lower left of the photo is interpreted as formed around a paramoudra burrow in the cretaceous sea floor. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 6 of 53 www.geusbul let in.org from marine algae (coccospheres). larger fossils of molluscs, bivalves, brachiopods, echinoids and siliceous sponges also occur. the chalk is commonly thoroughly bioturbated by benthic animals (surlyk 1971, 2017; surlyk & håkansson 1999). the chalk contains more or less coherent layers of black flint (fig. 3; madsen & stemmerik 2010), which locally provide evidence that the chalk was deposited on shallow banks (surlyk et al. 2006) in a sea that covered most of northern europe, and where clastic sedimentation was almost absent (fig. 4). however, thin cyclic deposits of marl occur in parts of the maastrichtien. the late cretaceous chalk in eastern denmark is divided into two formations: the mandehoved formation comprising two members (campanien) and the møns klint formation, which comprises five members and is named after the cliffs in the eastern part of møn (late campanien – maastrichtien; surlyk et al. 2013). the division is based on investigations at stevns klint (surlyk et al. 2006) and analyses from the c. 440 m stevns-1 deep borehole (fig. 5; surlyk et al. 2013). the type profile itself comprises several of the five members. the age of the chalk has been determined by means of macroand microfossils from the cliffs of høje møn and hvideklint, which yield ages from lower and upper maastrichtien (surlyk 1984; surlyk & håkansson 1999; thomsen 1995; jelby et al. 2014). glacial and interglacial deposits data and interpretation a packet of glacial sediments up to 50–100 m thick rests on the pre-quaternary surface. most of them were deposited during the last part of the weichselian glaciation c. 75  000–11  700 years ago. few traces of older deposits from the saalian glaciation (390  000– 130 000 years ago) and the eemian interglacial (130 000– 115 000 years ago) are also present. on most of møn, the glacial sediments have been deformed, folded and faulted by glaciotectonic processes, while large sheets of chalk have been pushed up from below the prequaternary surface under møn itself or surrounding areas, now covered by the sea (e.g. berthelsen 1979; pedersen 2000). along with slide-sheets of chalk in tills and coherent chalk particles in the clastic deposits the 50°n 60°n 0°ø 10°ø 20°ø 200 km møn n fig. 4 palaeogeographic map of the late cretaceous sea in northern europe. green: land areas. blue: sea. from gravesen et al. (2017) after damholt & surlyk (2012). age danian hvidskud mb boesdal mb flagbanke mb højerup mb sigerslev mb rørdal mb cerithium lmst mb fiskeler mbn e n eu pp er c re ta c eo u s m aa st ri ch ti an m øn s kl in t fm m an de ho ve d fm c am pa ni an lithostratigraphy fig. 5 stratigraphic division of the uppermost cretaceous and lower most danian in eastern denmark (after surlyk et al. 2006, 2013). http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 7 of 53 www.geusbul let in.org dislocated chalk sheets can be considered part of the quaternary deposits. during the mapping, the following glacial deposits were differentiated: clayey till (ml), sandy till (ms) and gravelly till (mg). add to these glaciofluvial gravel (dg), glaciofluvial sand (ds), glaciolacustrine silt (di) and glaciolacustrine clay (dl) as well as late-glacial deposits: freshwater gravel (fg), freshwater sand (fs), freshwater clay (fl) and chalk (sk). a complete overview of the deposits and symbols can be found in pedersen et al. (2015, appendiks 2). the quaternary stratigraphy is based on mapping, literature, data and studies of profiles on the southern, eastern and northern coasts of møn and in a few gravel pits (fig. 6a). in addition, the cross-sections shown on the geological map contribute much to the understanding of the glacial stratigraphy. a correlation at formation level using the cross-sections is shown in fig. 25. data sand clay silt fi ne c oa rs e g ra ve l c ob bl e bo ul de r clay & silt peat & gyttje sand sand, gravel & cobbles clayey till diamict flint glaciotectonite chalk laminated structureless ripples climbing ripples trough crossbedding planar cross-bedding sharp boundary clay with cobbles gradual boundary erosive or glaciotectonic boundary holocene eemian saalian chalk w ei ch se lia n holocene deposits lithology structures age 10 m late glacial freshwater layers lolland till formation ny borre formation mid danish till formation klintholm till formation kraneled formation ristinge klint till formation eemian deposits saalian clayey till maastrichtian chalk early weichselian deposits stubberup have formation kobbelgård formation lithology structures m ed iu m a fig. 6 a: schematic stratigraphic log of the quaternary interglacial, glacial and holocene deposits occurring in the møn map sheet area. b: locations of geographic names with reference to geological type localities. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 8 of 53 www.geusbul let in.org from geus’ borehole data archive and jupiter database have been used for the construction of these profiles (gravesen & fredericia 1984). a map of important geological localities on møn is shown in fig. 6b. older quaternary deposits: saalian, eemian, early and early middle weichselian (mis 6, 5e, 5a–5d and 4) a few scattered deposits of saalian and eemian age, marine isotope stage (mis) 6 age and even younger ages have been found in the cliffs (fig. 6a), forming fragments and small, thin sheets in the strongly deformed glaciotectonic units. despite being allochthonous, it is likely that deposits of this age have occurred in parts of møn. saalian (mis 6) a single layer of clayey till beneath eemian deposits at kraneled and two similar beds at both hellehavn nakke og stubberup have are assumed to be of saalian age (fig. 8; hyde 1986). several other localities probably also contain saalian clayey till, although the widespread glaciotectonic effects make this interpretation uncertain (puggaard 1851; hintze 1937; hyde 1986; houmarknielsen 1994). the clayey till layers are less than 3 m thick (fig. 7) and are probably of late saalian (warthe) age. the glaciotectonic deformation of both the saalian and eemian layers took place in the middle or late weichselian (berthelsen 1973; aber 1979; pedersen & gravesen 2009; houmark-nielsen 2010). eemian (mis 5e) in the eemian interglacial, denmark was surrounded by sea. it is assumed that the eemian sea covered parts of møn (seidenkrantz et al. 2000), although it is difficult to establish its precise extent because of the subsequent glaciotectonic disturbances. the presence of the eemian sea has been documented from the north sea, through the danish straits and far into the baltic sea in the east, presumably with an early eemian connection through karelen in finland to the white sea in the north-east (knudsen et al. 2011). the early eemian baltic transgression appears to have taken place from the east, with the oldest marine deposits in the finnish bay and the youngest in eastern denmark, the netherlands and germany (kristensen et al. 2000). the dominant deposits on møn are greenish to greenish black marine clay (cyprina clay). its thickness is locally up to c. 2 m (fig. 7) but in most places only about stubberup have hellehavns nakke liselund pomlerende taleren abildgaards fald dronningestolen gråryg siesø bjerg nælderende jættebrink kraneled klint kobbelgård klint tøvelde klint møn hjelm bugt stege bugt 12°30’ø hvid ek lin t stege hjelm nakke madses klint præstebjerg klintholm klint aborrebjerg ny borre 5 km 55°n nyord ulvshale kostervig ulvsund steg e n or b fig. 6 (continued) a: schematic stratigraphic log of the quaternary interglacial, glacial and holocene deposits occurring in the møn map sheet area. b: locations of geographic names with reference to geological type localities. weichselian clayey till, lolland till formation weichselian glacio�uvial sand weichselian clayey till, mid danish till formation weichselian, clay with minor clast content eemian marine clay saalian glacio�uvial sand saalian clayey till maastrichtian chalk sw ne 5 m ? ? ? a fig. 7 eemian deposits at hjelm nakke. a: field sketch of the saalian, eemian and weichselian deposits after berthelsen et al. (1977). b: clay and glaciofluvial sand (saalian?) overlain by marine eemian clay. photo: tove stockmarr (1996). http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 9 of 53 www.geusbul let in.org 0.5 m. the glaciotectonic deformations make these estimates uncertain. marine fossils have been found in sand and clay deposits along the north coast of møn at øksnehøj/ orebjerg, pomlerende and stubberup have and in several parts of møns klint, mainly in the falls but also along ridge flanks, e.g. at hundevangsfald, hvidskudsnakke, lille stejleberg, store stejlebjerg, grårygfald, fruerstuefald, nælderende, sandfaldet, skrædderrenden, magle vandsklinten/faldet, dronningestolen, vitmunds nakke, søndre hylledalsklint, taleren and jydeleje fald (puggaard 1851; madsen et al. 1908; ødum 1933; hintze 1937; hansen & nielsen 1960; hyde 1986). eemian deposits also occur at kraneled klint east of klintholm on the south coast of møn (houmark-nielsen 1994) and at two localities in the cliff between hjelm nakke and hvideklint (fig. 7; berthelsen et al. 1977; stockmarr 1996). in addition, gyttja and sand with freshwater molluscs and plant fossils from an interglacial period have been found under marine deposits, for example at store stejlebjerg (n. hartz in hintze 1937). this eemian transition from freshwater to marine deposits is also known from other baltic localities (kristensen et al. 2000; knudsen et al. 2011). macrofossils have been found in both sand and sticky clay. marine molluscs in the clay have commonly been affected by glaciotectonic deformation, but nevertheless appear to remain in situ. the number of different species (ødum 1933; hintze 1937) is small compared to the extensive list of marine eemian fossils listed in madsen et al. (1908). the lithology, age and fossil content of the fossiliferous deposits have been debated, but foraminiferal studies have demonstrated marine eemian faunas at hjelm bugt, kraneled and pomlerende (v. madsen in hintze 1937; berthelsen et al. 1977; j. frederiksen in hyde 1986; p. kristensen in houmark-nielsen 1994); this supports the presence of marine macrofossils in these deposits. local occurrences of sand and clay intercalated between layers of clayey till, which all contain eemian foraminifera, show that these occurrences have been redeposited and that the beds are younger than eemian (konradi 1973). the foraminifera in the eemian deposits belong to a boreo-lusitanian, nearshore shallow-water fauna (berthelsen et al. 1977), comparable to faunas from other localities in the western baltic sea and in the north sea (konradi 1976; knudsen 1991). these species lived in sea water that was a little warmer and more saline than at present (houmark-nielsen 1994). age determinations of eemian deposits from klintholm–kraneled by means of optically stimulated luminescence (osl) and thermoluminescence (tl) have yielded considerably younger ages than expected (younger than 115 000 years before present (bp)). this discrepancy was discussed by houmark-nielsen (1994, 2008, 2010). amino acid dating of selected molluscs (arctica islandica, turritella communis and nassa reticulata from klintholm, arctica islandica and turritella communis from stubberup have and turritella communis from hjelm nakke (petersen 1984; miller & mangerud 1985; houmark-nielsen 1994, 2008)) supports an eemian age of the deposits from these cliff profiles. early and early middle weichselian (mis 5a5d, mis 4) in early and early middle weichselian, møn was probably ice-free, with surface conditions dominated by solifluction and limited fluvial transport and deposition. redeposited eemian macrofossils occur in deposits from this period (houmark-nielsen 1994), which have estimated thicknesses of 2–20 m. fig. 7 (continued) eemian deposits at hjelm nakke. a: field sketch of the saalian, eemian and weichselian deposits after berthelsen et al. (1977). b: clay and glaciofluvial sand (saalian?) overlain by marine eemian clay. photo: tove stockmarr (1996). b http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 10 of 53 www.geusbul let in.org middle and late weichselian (mis 3, 2) the most widespread glacial units in the map area were deposited in middle weichselian (74 000–29 000 years ago) during the last part of the latest ice age and in late weichselian (29 000 to c. 11 700 years ago, fig. 6a), but only parts of these time intervals are represented. because of the extensive glaciotectonic disturbance of the quaternary deposits on møn, it is difficult to make appropriate formal descriptions of existing lithostratigraphic units, establish new ones and document their extent, despite good exposure. nevertheless, a number of geological cliff profiles with descriptions of the lithologies and interpretations of their formation have been published in recent years, for example by houmarknielsen (1994) and gravesen et al. (2017) in reports and unpublished msc theses. in addition, hydrogeological mapping (rambøll 2006, 2007) has contributed to the present understanding of the three-dimensional geological structure in the interior parts of møn where there are few outcrops. this information, combined with the systematic surface mapping, profiles in gravel pits, information from boreholes and geophysical data, has contributed to the establishment of new and revisions of existing middle and late weichselian geological formations presented in the following. all quoted ages are based on houmark-nielsen et al. (2017). a profile along southern, eastern and north-eastern møn has been constructed from selected logs (fig. 8). the profile begins at madses klint in the south, follows the south-eastern and eastern coasts with møns klint, continues towards north-east to hellehavn nakke and terminates on the north coast at brunshoved. the profile illustrates the overall lithostratigrafic structure and the horizontal and vertical extent of each unit. the different units and their formal division into formations are described below. ristinge klint till formation (mis 3) name and history. the formation is named after ristinge klint on langeland, a classical locality which has contributed to the general understanding of the geological structure of denmark (madsen et al. 1908; lolland till fm kobbelgård fm klintholm till fm kraneled fm ristinge klint till fm early weichselian saalian clayey till late glacial freshwater layer ny borre fm mid danish till f m stubberup have fm 0 60 km m ad se s k lin t hjelm n ak ke tøve lde klin t north ern m øns k lin t helle hav ns n ak ke stu bberu p h av e bru nsh av pomler en de klin th olm lis elund hvid e klin t 1 0 20 28 32 33 36 39 5 1 0 1 5 20 25 30 36 west east west stubberup have fm mid danish till fm ny borre fm lolland till fm m north-east w e w ne eemian deposits fig. 8 composite section showing the thicknesses and horizontal distribution of the lithostratigraphical units along the south, east and north coast cliffs of møn based on selected vertical logs. the index map shows where the section changes orientation. note different horizontal and vertical scales. based on houmark-nielsen (2010) and gravesen et al. (2017). http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 11 of 53 www.geusbul let in.org madsen 1916). andersen (1945) and berthelsen (1973) called the till unit the “gammel baltiske till” (old baltic till). houmark-nielsen (1987) formally erected the ristinge klint till formation and presented a detailed description of its history of formation. he showed that the deposition of the till, which contains fine-grained gravel with a baltic provenance, was related to an ice advance from south-east (sjørring et al. 1982). on møn, houmark-nielsen (1994) called the formation “grey-red till, unit 4” at klintholm. type locality. ristinge klint on langeland (houmarknielsen 1987). lithology. the formation consists of grey to reddish grey till. the reddish colour is due to a high content of lower palaeozoic baltic sandstone and limestone. the lower part of the formation consists commonly of layered, grey to whitish till with a high content of fragmented, sheared and folded chalk and smeared out by glaciotectonic deformation, a glacitectonite (fig. 9; banham 1977; pedersen 1988). the upper part of the till is reddish grey with reddish sand and sandstone, which can be commonly found on the beach below the cliffs. fine-grained gravel and clasts are dominated by baltic material (berthelsen et al. 1977; hyde 1986; houmark-nielsen 1994; stockmarr 1996; smed 2010). redeposited eemian material occurs locally. in the cliff profiles of north-eastern møn, till fabric measurements show a poorly defined easterly orientation with a shallow to moderate easterly plunge (50 measurements in hyde 1986). disregarding the fact that the till has also been affected by subsequent glaciotectonic dislocation in this area, these measurements show a well-defined ice movement from east to west. this provenance is supported by indicator counts of fine gravel showing 25% limestone (of which two thirds are palaeozoic), a low content of flint (<5%) and up to 15% sedimentary gravel, mainly consisting of sandstone (hyde 1986). lithological symbols: ml, ms. boundaries and thickness. at klintholm, the lower boundary of the formation is a glaciotectonic un conformity above older, presumably early or early middle weichselian deposits. at hvideklint and hjelm nakke, the formation commonly rests on chalk. the upper boundary against the kraneled formation is sedimentary and erosive. the thickness of the formation is probably a few metres at hvideklint – hjelm nakke, 4 m at klinthavn, 1 m at høje møn, 3 m at hellehavn nakke and 2 m at pomlerende (hyde 1986; houmark-nielsen 1994; stockmarr 1996). extent and correlation. the formation occurs at a number of different localities, for instance hjelm nakke – hvideklint (probable), høje møn, klintholm, hellehavn nakke and pomlerende. this indicates that the formation previously had a larger extent in southern and eastern møn, and it probably exists in large parts of denmark (houmark-nielsen 1987, 2007). age. the formation was deposited as part of the ristinge ice advance between 55  000 and 50  000 years ago (houmark-nielsen 2010; houmark-nielsen et al. 2017). depositional environment. the till is a lodgement till. its content of baltic fine-grained gravel and clasts and the glaciotectonic structures show that the glacier movement was from south-east towards north-west mid danish till fm kobbelgård fm risti nge klin t t ill f m cha lk fig. 9 red-brown clayey till of the ristinge klint till formation at store stejlebjerg, møns klint. the formation rests unconformably on top of the cretaceous chalk of the møns klint formation. fine-grained lithologies of the kobbelgård formation overlie the ris tinge klint till formation. at the top, clayey till of the mid danish till formation is present. for scale, the units on the ranging pole are 20 cm wide. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 12 of 53 www.geusbul let in.org (houmark-nielsen 1999; smed 2010), called the ristinge ice advance (houmark-nielsen 1987; houmark-nielsen et al. 2017). kraneled formation (mis 3) new formation name and history. the kraneled formation is named after the locality of kraneled east of klintholm on southern møn (fig. 6b). houmark-nielsen (1994) described it as “glaciofluvial and lacustrine deposits, unit 5”, while it was called “klintholm beds” in houmark-nielsen & kjær (2003) and “klintholm” in houmark-nielsen et al. (2017). type locality and type profile. the type locality is kraneled klint east of klintholm (fig. 10; houmark-nielsen 1988, 1994). the type profile is shown in houmark-nielsen (1994, fig. 5, unit 5). lithology. the lower part of the formation consists of cross-bedded and imbricated sand and gravel. the upper part is cross-bedded sand changing upwards into fine-grained sand and horizontally laminated clay (fig. 10). a few plant remains may be present. lithological symbols: dl, ds, dg. boundaries and thickness. the lower boundary against the ristinge klint till formation is sharp and erosional. the upper boundary is a gradual sedimentary transition to the klintholm till formation. the thickness of the formation is around 4 m at klintholm and 2 m at tøvelde klint, but larger in western møn, about 7 m in wells. extent and correlation. because of poor exposure, it is difficult to establish the lateral extent of the formation. it occurs along hjelm bugt and on western møn. drilling investigations at kriegers flak in the baltic sea east of møn document the presence of interstadial lacustrine deposits 40  000–35  000 years ago, which probably correspond to the kraneled formation (anjar et al. 2010, 2012). deposits of the same age occur within kraneled fm ristinge klint till fm limestone glaciotectonite chalk fig. 10 fine-grained meltwater deposits of the kraneled formation resting on the ristinge klint till formation. between the till and the underlying chalk, a chalk-glacitectonite occurs. kraneled cliff, klintholm. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 13 of 53 www.geusbul let in.org the large ice-dammed lake area at brorfelde on sjælland, where mammoth tusks have been found in thick sandy deposits dated at 44 000 bp (selsing 1982). interstadial deposits on sejerø west of sjælland have been dated to c. 36 000 years ago, i.e. prior to the deposition of the klintholm till formation and probably contemporaneous with the kraneled formation deposits. the plant content in the lacustrine deposits on sejerø indicate an open, treeless area with tundra vegetation (kolstrup & houmark-nielsen 1991; bennike et al. 2007). the fine-grained part of the kraneled formation may represent the same depositional environment as on møn during the sandnæs interstadial period. age. the formation has an age of c. 45  000–35  000 bp (houmark-nielsen et al. 2017) and was therefore probably formed during the first phase of the sandnæs interstadial. depositional environment. the lowermost, coarsegrained glaciofluvial material was probably deposited in a meltwater river in a periglacial area. the overlying deposits point to lacustrine deposition. the landscape was presumably open, with sparse arctic tundra vegetation. klintholm till formation (mis 3) revised formation name and history. the formation is named after the small town of klintholm on the south coast of møn. it has been described by houmark-nielsen (1994) as “thick grey diamicton” and by krüger og kjær (1999) as “unit 6”. the till deposit has subsequently been described as “klintholm till” (houmark-nielsen & kjær 2003; houmark-nielsen 2008, 2010; houmark-nielsen et al. 2016). type locality and type profile. kraneled east of klintholm. a reference profile is found at kobbelgård west of klintholm (fig. 11; houmark-nielsen 1988, 1994, fig. 5, unit 6). lithology. the klintholm till formation consists of massive, silty and sandy till with a moderate clast content (figs 6, 11). in its uppermost part dark brown, massive and locally stratified clayey till may occur. redeposited shells and shell fragments of marine eemian molluscs may also occur. lithological symbols: ml, ms. boundaries and thickness. the thickness of the formation is c. 2.6 m at madses klint, 3.6 m at hjelm nakke – hvideklint, 4 m at tøvelde klint, 3–5 m at klintholm, 3 m at møns klint, c. 1.5 m at hellehavn nakke and 1.8 m at stubberup have (hyde 1986; houmark-nielsen 1994; stockmarr 1996; gravesen et al. 2017). the lower boundary of the klintholm till formation is mostly a gradual transition from clayey till downwards to sand and mud of the kraneled formation, but locally, may also be a glaciotectonic unconformity. the upper boundary is gradational from clayey till to mud of the kobbelgård formation. extent and correlation. the klintholm till formation has a wide extent from møn to falster (korselitse) and southern sjælland (houmark-nielsen 2010). age. the klintholm till formation was laid down during the klintholm ice advance, dated at c. 35 000–32 000 bp (houmark-nielsen 2010; houmark-nielsen et al. 2017). depositional environment. the formation is interpreted as a lodgement till deposited by a glacier with a general movement from south-east towards north-west. this can be demonstrated by the content of indicator erratics of fine gravel, clast fabric orientations and glaciotectonic structures. smed (2010) counted and analysed clasts and showed that a part of the glacier movement was first from north-east to south-west through sweden, and then changed to a westerly direction towards denmark (see also ukkonen et al. 2007). the ice advance was called the “klintholm isstrøm” by houmark-nielsen et al. (2017). kobbelgård formation (mis 3) new formation name and history. the kobbelgård formation is named after the cliff at kobbelgård, hjelm bugt, southern møn, where interstadial lacustrine, glacial lacustrine and glacial fluvial deposits occur. the deposits have been studied particularly well near klintholm havn in the two cliffs at kobbelgård and kraneled (houmark-nielsen 1988, 1994, 2003, 2010; krüger & kjær 1999; houmarknielsen et al. 2016), where a sequence that also comprises the klintholm till and kraneled formations (fig. 8) has been described in detail and interpreted based on age determinations, sedimentology and the content of plant and animal fossils. the lower part of the clayey deposits is known from previous investigations on møn (e.g. puggaard 1851; hintze 1937; hansen & nielsen 1960; berthelsen et al. 1977) and described as a clay deposit with minor clast content (“stenfattigt ler”, hyde 1986). a characteristic vertical joint pattern (“søjleforkløftning”, berthelsen et al. 1977) has led to the informal name elephant-foot clay because eroded and commonly folded exposures may http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 14 of 53 www.geusbul let in.org resemble elephant feet. at klint holm, the layers have been described by houmark-nielsen (1994) as “lower lake sediments, unit 7” as a part of “upper kobbelgård beds” by houmark-nielsen et al. (2016). the upper part of the formation mainly comprises the fossil-rich part of the lacustrine deposits at klintholm, which consist of clay, silt and diamict deposits. they have been described as “upper lake sediments, unit 8”, by houmark-nielsen (1994) and “upper kobbelgård beds” or “kobbelgård beds” (houmark-nielsen & kjær 2003; houmark-nielsen 2010). they have also been described by krüger & kjær (1999). type locality and type profile. the type locality for the kobbelgård formation is kobbelgård klint and kraneled klint, west and east of klintholm havn, respectively, where several partial profiles have been measured (see houmark-nielsen 1994 fig. 5, unit 7, unit 8; houmarknielsen 1988, 1994; krüger & kjær 1999). lithology. the lowest part of the formation consists of clay/mud with thin sandy layers and many dispersed clasts (fig. 12). the clayey deposit contains numerous joints. cross-bedded and horizontally bedded sand occurs locally at the base of the formation, in places with redeposited shells derived from eemian deposits. the clay is sticky and silty, calcareous and dark grey, with thin layers of sand and commonly dispersed clasts. the content of clay and silt is 90–94%; the remainder is sand and fine-grained gravel, the latter almost entirely consisting of palaeozoic limestone (hyde 1986). apart from joint systems, the only visible structure is indistinct layering in the clayey part, which is difficult to discern. scattered plant remains consist of leaves and detrital plant material, as well as pollen from an arctic flora (kolstrup & houmark-nielsen 1991). the upper part of the formation consists of a number of alternating layers of massive clay and silt as well as discontinuous layers of fine sand with intercalations of clayey diamictites towards the top (fig. 13). plant remains including leaves and fine-grained detritus also occur, and an extensive microog macroflora has been described (kolstrup & houmark-nielsen 1991; bennike et al. 1994, 2007). clasts and fine gravel in the diamict deposits are dominated by palaeozoic limestone and shale, but the uppermost layers contain a mixture of local material and rock fragments from the baltic area. lithological symbols: dl, di, ds, ml. boundaries and thickness. the lower boundary towards the klintholm till formation is commonly gradational. the upper boundary to the mid danish till formation is a glaciotectonic unconformity. the thickness of the kobbelgård formation reaches up to 15 m at klintholm but varies at other localities, with approximate values as follows: madses klint: 7.5 m, hjelm nakke – hvideklint: 12.5 m, tøvelde klint: 7.5 m, møns klint: 12 m, liselund: 5 m, hellehavn nakke: 11 m and stubberup have: 7 m (ødum 1933; hintze 1937; hyde 1986; houmark-nielsen 1988, 1994; stockmarr 1996; houmark-nielsen et al. 2016; gravesen et al. 2017). extent and correlation. the formation has a wide extent on southern, eastern and north-eastern møn. the upper part, rich in fossils, is mainly encountered on the south coast. parts of the formation are probably contemporaneous with deposits on kriegers flak (anjar et al. 2010, 2012). the formation is also exposed along the east coast of falster, for example at korselitse (houmark-nielsen 2010; gravesen et al. 2017). age. the deposits are from the middle weichselian, mis 3. the formation rests on the klintholm till formation, which has an age of 35 000–32 000 bp (houmark-nielsen et al. 2017). lolland till fm kobbelgård fm klintholm till fm ristinge klint till fm chalk fig. 11 grey clayey till of the klintholm till formation at kraneled cliff, klintholm. the section also contains deposits of chalk as well as deposits of the ristinge klint till, kobbelgård and lolland till formations. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 15 of 53 www.geusbul let in.org a number of 14c, osl and tl age determinations on plant material and sediments from the kobbelgård formation have yielded ages spreading through the last part of the middle weichselian, with ages of 35  000– 20 000 years using the osl method and 37 000–25 000 using the 14c method (houmark-nielsen 2010). collectively, the age determinations from klintholm indicate that the most likely age is between 32 000 and 28  000 bp (houmark-nielsen & kjær 2003; houmarknielsen 2008, 2010; houmark-nielsen et al. 2016, 2017). the age of the formation is 32  000–28  000 bp, representing deposition during the last part of the sandnæs interstadal (sandnæs mildningen, houmarknielsen et al. 2017). depositional environment. at the end of the cold period, about 32 000 years ago, when the ice of the klintholm ice advance melted away, a large glacial lake was formed, which covered all of southern, eastern and northern møn and possibly also an even larger baltic area, as suggested by hansen (1965) and houmark-nielsen (1994). sand and gravel were deposited first, where meltwater entered the lake. widespread deposition of clay and silt followed, and ice rafts transported clasts a b fig. 12 lacustrine deposits of grey, silty clay with scattered pebbles of the kobbelgård formation. a: typical glaciolacustrine clay with widely spaced dropstones. b: clay at liselund in the dislocation horizon at the toe of a landslide. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 16 of 53 www.geusbul let in.org into the lake, which were deposited as dropstones. fine sand was blown out over the lake from the barren surrounding landscape and deposited as thin strata. as the climate became warmer and less meltwater was supplied from the baltic ice, the size of the lake diminished, and an arctic vegetation with grasses herbs and a few small trees like pinus and betula was established around the lake. from 30  000 to 28  000 bp, the lake was probably restricted to an area around klintholm. mud (clay, silt and fine sand) was deposited from suspension, while diamictites were deposited from mud streams and/or ice rafts. also, fine aeolian sand and silt deposits have been documented (houmark-nielsen 1994; krüger & kjær 1999). the vegetation around the lake was subarctic with many herbs and shrubs, but treeless. the increasingly warmer subarctic climate led to increasing vegetation in and around the lake. lateral cross-bedded, fine-grained dunes were laid down by adjacent river systems, which also carried sand and gravel out into the lake (the stubberup have formation). a mammoth tooth found at slotshøj, stege nor (aarissørensen et al. 1990) may indicate the former presence of a large grassy area around the lake. at this time, large grassy and herbaceous plains with mammoths and other large animals have been present in denmark (aaris-sørensen et al. 1990) and southern sweden (berglund et al. 1976; ukkonen et al. 2007). stubberup have formation (mis 3) new formation name and history. named after the northern part of møns klint in the area around stubberup have. type locality and type profile. stubberup have (fig. 14; hyde 1986). lithology. the formation consists of sand, gravel and occasionally stones (figs 8, 14). sedimentary structures comprise ripples, horizontal bedding, large-scale fig. 13 grey, silty and fine-grained, horizontally laminated clay of the kobbelgård formation. kobbelgård cliff, klintholm. for scale, the units on the ranging pole are 20 cm wide. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 17 of 53 www.geusbul let in.org cross-bedding and coarse-grained channel structures with gravel and stones, which have incised into the underlying layers of sand. fine-grained, well sorted sand and silt with climbing ripples are also encountered. boundaries and thickness. the thickness of the stubberup have formation is quite variable (fig. 8), with the following observations: madses klint: 1.3 m, hjelm nakke – hvideklint: 4.5 m, tøvelde klint: 5 m, møns klint: 1.4 m, liselund: 10 m, hellehavn nakke: 11 m and pomlerende: at least 3 m (hintze 1937; hyde 1986; stockmarr 1996; gravesen et al. 2017). a few boreholes point to thicknesses of up to 20 m, but the layers may have been disturbed by glaciotectonics. the formation has a sharp lower depositional boundary against the kobbelgård formation (fig. 8). towards west from the tøvelde and klintholm cliffs towards madses klint, the boundary is probably interdigitating. the upper boundary against the mid danish till formation is a sharp glaciotectonic unconformity. a clay & silt sand sand sand, gravel & cobbles lithology laminated cross-bedding structures sharp boundary ? clayey till clay silt �n e m ed iu m co ar se g ra ve l & co bb le s 2 m 1 2 3 0 4 5 8 7 6 11 10 9 12 mid danish till formation stubberup have formation kobbelgård formation b fig. 14 a: mediumand coarse-grained, cross-bedded sand of the stubberup have formation at the cliff of stubberup have, the type locality for the formation. the units on the ranging pole are 20 cm wide. b: sedimentological log of the stubberup have formation. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 18 of 53 www.geusbul let in.org extent and correlation. the stubberup have formation is found in large parts of møn. age. the age of the stubberup have formation is likely the same age as the kobbelgård formation (32  000– 28 000 bp; houmark-nielsen et al. 2017). depositional environment. the depositional environment was a plain with braided rivers, which formed small ripples and larger dunes. lacustrine conditions locally prevailed, and the deposition of silt and fine-grained sand as climbing ripples suggests a large sediment supply. mid danish till formation (mis 2) name and history. the mid danish till formation was established by houmark-nielsen (1987) for the lithological till unit that extends from the eastern boundary of denmark to the eastern jutland stationary line. at the type locality of ristinge klint on langeland, the formation forms a homogenous lithological unit about 15 m thick within the imbricated sheets that comprise the glaciotectonic complex of ristinge klint. the geological basis for the establishment of the formation can be found in houmark-nielsen (1987). fig. 15 grey clayey till of the mid danish till formation, madses klint. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 19 of 53 www.geusbul let in.org type locality and type profile. ristinge klint on langeland (houmark-nielsen 1987). the reference profile on møn is at madses klint (fig. 15). lithology. the formation consists of massive, calcareous till, commonly with a high content of chalk clasts. where the till has been affected by glaciotectonic shear, the chalk erratics may be classified as “glacitectonite” (pedersen et al. 2018). lithological symbol: ml. boundaries and thickness. the formation has a sharp contact against the underlying stubberup have formation, usually a glaciotectonic unconformity (fig. 6). it is overlain by the ny borre formation with a sharp depositional boundary. on møn, the thickness of the mid danish till formation is very variable but may reach a thickness of up to 10 m. the following thicknesses have been observed: madses klint: 3.5 m, hjelm nakke – hvideklint: 6 m, tøvelde: 4 m, møns klint: 9 m, liselund: 3 m, pomlerende: variable thicknesses, stubberup have: 5.5 m and brunshoved: 8 m. (ødum 1933; hansen & nielsen 1960; berthelsen et al. 1977; hyde 1986; stockmarr 1996; gravesen et al. 2017). extent and correlation. the mid danish till formation covers all of møn except the klintholm area. age. the age of the formation is c. 23  000–20  000 bp (houmark-nielsen et al. 2017) and it was deposited during the late glacial maximum (lgm) of the weichselian glaciation. depositional environment. the mid danish till formation is a lodgement till deposited by a glacier. glaciotectonic orientation parameters, clast content and indicator counts of fine gravel show that the glacier moved from north-east towards south-west (berthelsen et al. 1977; aber 1979; hyde 1986; stockmarr 1996; smed 2010). the formation was deposited during the maximal extent of the scandinavian ice cap in the last part of the weichselian glaciation. the main stationary line for this ice advance was located 100 km south and west of møn, stretching from flensburg and north of hamburg to brandenburg and then turning south of berlin and into northern poland, the so-called maximum brandenburg ice-marginal stationary line (smed & ehlers 2002; ehlers et al. 2004). ny borre formation (mis 2) new formation name and history. the formation is named after the village of ny borre south of borre on eastern møn. it includes the surficial glaciofluvial and glaciolacustrine deposits of sand and gravel that occur in large parts of møn. these deposits have not previously been formally named. sand and gravel from the ny borre formation are excavated, for example, in the ny borre gravel pit. type locality and type profile. the type locality is in the active gravel pit at ny borre. the type profile is shown in fig. 16, a layered sandy sequence comprising both coarseand fine-grained deposits. lithology. the formation consists of layered silt, sand and gravel. the coarse-grained sand is found in many places on møn. the grain size is variable, depending on the local depositional environment. the lower part consists of alternating layers of fineto medium-grained sand and silt as well as lenses of coarse-grained sand and occasional layers of clay. the layers are inclined locally and may be part of channel deposits of larger foresets, which point to a current direction from south and southwest (fig. 17). the upper and more widespread part consists of cross-bedded, coarse sand, intercalated with gravel in larger channel structures. lithological symbols: ds, dg, di, fs, fg. clay & silt sand sand lithology laminated structures sand, gravel & cobbles ripples trough cross-bedding sharp boundary structureless clayey till clay silt �n e m ed iu m co ar se g ra ve l & co bb le s 2 m fig. 16 sedimentological log of the fluvial–lacustrine sequence of the ny borre formation at the type locality in the ny borre gravel pit. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 20 of 53 www.geusbul let in.org boundaries and thickness. the lower formation boundary against the mid danish till formation is sharp and mainly depositional. the lower and upper boundaries can be observed at the same localities. the upper boundary against the lolland till formation is also sharp and commonly a glaciotectonic unconformity. it can be observed in a number of coastal cliffs, for example madses klint, hjelm nakke, hvideklint, tøvelde, møns klint and in the cliffs north of høje møn (liselund, stubberup have and pomlerende). the thickness of the formation is 5–10 m in large parts of the island but may be greater in the large channels. at the type locality of the ny borre gravel pit, the thickness is c. 6 m (storstrøms amt 2004; region sjælland 2012). the thickness variation in the cliffs, where the formation is overlain by the lolland till formation, is as follows: madses klint: 12 m, hvideklint: 6 m, tøvelde: 1.5 m, møns klint: 5.5 m, liselund: at least 4.5 m, pomlerende: up to 3 m and stubberup have: 5.5 m (hintze 1937; hyde 1986; stockmarr 1996). extent and correlation. the formation extends over large parts of møn and is shown on the geological map as a surficial deposit or immediately under the lolland till fig. 17 inclined, alternating silt, sand and gravel layers in a lacustrine sequence of the ny borre formation. ny borre gravel pit. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 21 of 53 www.geusbul let in.org formation. this is also observed in rambøll (2006). here, its extent follows a number of valleys, from north to south comprising borre sømose, maglemose, klostervig and the lowland bordering stege nor, the røddinge lowland and fanefjord, which are all tunnel valleys or draining channels from the melting young baltic ice. the esker of råbylille ås is located in a supposed tunnel valley extending from kobbelgård to stege nor. the esker of fanefjord ås is found in the fanefjord tunnel valley (smed 2014). a variety of the ny borre formation sediments is found in both of these tunnel valleys, as well as in hat-shaped hills in south-western møn (berthelsen 1980). age. the ny borre formation is regarded as having been deposited during the young baltic ice advance, presumably 19 000–18 000 bp (houmark-nielsen et al. 2017). depositional environment. the formation represents several successive depositional environments formed by meltwater from the young baltic ice: a glaciofluvial environment where sand and gravel were deposited on valley sandurs, which developed in a number of large draining channels eroded into the mid danish till formation, and a glaciolacustrine environment with deposition in variably ice-dammed elongate lakes perpendicular to the ice margin. layers of sand and gravel were also deposited in near-surface eskers and hat-shaped hills. lolland till formation (mis 2) name and history. the lolland till formation is named after lolland in south-eastern denmark and was established by pedersen et al. (2015) in the sakskøbing map sheet description. the till that covers the landscapes of south-eastern denmark are traditionally assigned to the lodgement till deposited by the young baltic ice advance. in the southern and central parts of eastern jylland, the till is called the east jylland till formation (houmark-nielsen 1987), while lodgements from the subsequent glacial re-advance, which completed the ice cover of denmark, are called the bælthav till formation. type locality and reference profile. the type locality of the lolland till formation is the birket gravel pit on lolland (pedersen et al. 2015). on møn, there is a reference profile at kobbelgård (fig. 18; houmark-nielsen 1994). together with the ny borre formation, the lolland till formation forms the predominant surficial lithology of møn, comprising 80% of the land area. access to exposures through the formation can be difficult as the till forms the uppermost part of the relatively high coastal profiles. here the formation is seen to be draped over all other lithologies, and in several of the profiles it rests unconformably on older deposits, disturbed by glaciotectonic processes. lithology. the formation consists of silty, brown to yellowish brown till (fig. 19). the lithology is commonly massive and compact but may also contain a weak horizontal stratification. the clast content is relatively low, although a few clasts of chalk are present (houmark-nielsen 1994; smed 2010; gravesen et al. 2017). till layers with a very high content of chalk may be classified as chalk-glacitectonite (banham 1977; pedersen 1988), as described in pedersen et al. (2018). counts of indicator erratics of fine gravel from several localities (hyde 1986; houmark-nielsen 1994) show that palaeozoic carbonate grains are more common than cretaceous ones. lithological symbols: ml, ms. boundaries and thickness. the lower boundary is the glaciotectonic unconformity at the base of the formation. the formation mostly rests on the ny borre formation, except in the klintholm area where it rests on the kobbelgård formation (fig. 8). in large parts of møn, the upper boundary is also the terrain surface. however, it is overlain by late weichselian freshwater deposits, for example at tøvelde klint and hjelm nakke, and locally by holocene marine deposits. in both cases the boundaries are sharp. the thickness is generally 1–3 m. borehole data suggest it may be thicker, although potentially affected by glacial tectonic disturbance. in the cliffs, the following thicknesses have been measured: madses klint: 2 m, hvideklint: 2 m, tøvelde klint: 2.8 m, klintholm: 0.8 m, northern møns klint: 0.8 m, stubberup have: 2 m, liselund: 1.5 m, pomlerende; 0–2 m, hellehavn nakke: 2 m and brunshoved: at least 2 m (hyde 1986; gravesen et al. 2017). thin strata from the formation have also been found in the hat-shaped hill of præstebjerg on southwestern møn (berthelsen 1980). the formation is absent from the highest part of møns klint. extent and correlation. the lolland till formation extends over the entire area of lolland, falster and møn and on the islands of the strait of smålandsfarvandet. age. the lolland till formation has an age of 18  000– 17 000 bp (noe-nygaard et al. 2017). depositional environment. the melting of the northeastern ice left møn, falster, lolland and the strait of smålandsfarvandet as a hummocky moraine plateau. this was subsequently covered by the young baltic ice http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 22 of 53 www.geusbul let in.org advance, which spread from the north-eastern baltic sea towards west. analyses of clast fabrics and glaciotectonic deformation combined with a high content of palaeozoic limestone show that the till was deposited by a glacier moving from south-east towards north-west (hyde 1986). while the young baltic ice was advancing, its frontal part was spread out with movements towards its margins, with an advance towards central denmark on its northern flank and towards northern germany on its southern flank. on lolland, this meant that the initially e–w-oriented ice stream was turned counterclockwise into a nw-directed flow. finally, the flow direction was rotated still further towards nnw the nwand nnwdirected ice advance stopped at the eastern jutland stationary line in the central part of djursland, where the meltback resulted in the extramarginal outwash plain of tirstrup hedeslette (pedersen & petersen 1997). from here, the ice melted back with oscillatory movements, which resulted in the bælthav ice advances, where a series of stationary lines have been mapped (milthers 1948). on møn, this resulted in till deposition, while the lolland till formation was split into two till units separated by a thin layer of meltwater sand in parts of lolland and falster. late weichselian (late glacial, mis 2) deposits from the last part of late weichselian (senglacial) are mainly known from hjelm and tøvelde on the south coast of møn, where freshwater deposits comprise large parts of the late glacial period between 17 000 and 11 700 bp: oldest dryas, bølling, older dryas, allerød and younger dryas (kolstrup 1982; noe-nygaard & heiberg 2001). tøvelde, which is one of the best investigated localities, comprises an almost complete late glacial profile (fig. 20). late glacial freshwater deposits have also been found in bogs on høje møn (hintze 1937), whereas no marine deposits are known from this period on møn. at madses klint, there are layers of calcareous gyttja and peat on top of the young baltic lolland till formation. these lolland till fm fig. 18 greyish brown clayey till of the lolland till formation in the uppermost part of the cliff at kraneled, klintholm. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 23 of 53 www.geusbul let in.org layers may be either late glacial or postglacial in age but have not been dated (fig. 21). late glacial deposits also occur at the base of the borre sømose profile (mikkelsen 1949). the freshwater sediments mainly consist of clay (tl), silt (ti), gyttja (tp), peat (tt), sand (ts), gravel (tg) and variegated layers (tv). the deposits at tøvelde are well exposed along c. 100 m in the coastal cliff at tøveldestenen (the tøvelde boulder) and comprise several small lake basins, which gradually developed into a larger lake. the exposure at hjelm is more restricted, located high up in the coastal profile and difficult to access (heiberg 1991). the lake basins of these two well-described localities show an extensive freshwater fauna and flora as well as terrestrial animals from around the lakes (johansen 1904; heiberg og bennike 1997; noe-nygaard & heiberg 2001). the fauna comprises terrestrial rodents, amphibians, snails, bivalves and insects as well as freshwater fish, snails, bivalves and insects. the deposits are collectively up to 3.8 m thick. deposits in several bogs on høje møn have also been studied (hintze 1926, k. jessen in hintze 1937). in some of them, a late weichselian freshwater sequence of clay from older dryas, gyttja from allerød and clay from the younger dryas have been found, overlain by holocene deposits. late weichselian development (late glacial) in the late weichselian, at about 16 000 bp, a large icedammed baltic meltwater lake, the baltic ice lake was formed. on møn, a variable deposition of mainly peat and gyttja took place in the smaller lake depressions at hjelm og tøvelde. the lake basins were formed in an elevated area by the melting of isolated ice bodies left behind from the retreating young baltic ice. the climatic conditions constituted an important control on the sedimentation in the shallow lakes, but the deposition was also affected by variations in the elevation of the baltic ice lake (kolstrup 1982; noe-nygaard & heiberg 2001; noe-nygard et al. 2017; rosentau et al. 2017). in the oldest dryas, a low tundra vegetation spread over a previously barren landscape. sand, silt and resedimented till with plant material was washed out into the shallow lakes at hjelm and tøvelde or transported from the bordering areas by solifluction. fig. 19 grey-brown, clayey till of the lolland till formation below late weichselian freshwater deposits at tøvelde klint. at the top of the lacustrine succession a calcareous gyttja is seen. fig. 20 late weichselian – holocene lacustrine deposits at tøvelde, formed in a shallow depression of the till of the young baltic ice advance. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 24 of 53 www.geusbul let in.org the plant material comprises for example salix herbacea, artemisia, poaceae and cyperaceae. in the warmer bølling interstadial, plants such as sea buckthorn and dwarf birch were introduced along with dryas and horsetail. this led to an increasing content of humus in the soil. increasing precipitation caused leaching of chalk from the calcareous till and sand into newly established shallow lakes. in these lakes, fossils of fish have been encountered, and land snails lived around them. in the last part of the bølling interstadial, sand with a humus content was deposited in a boggy, lacustrine environment. the climate became colder and dryer during the older dryas, and dwarf birch declined. calcareous clay with dropstones was deposited in the lakes. unsorted soil was washed out into the basins, and the leaching of calcareous material stopped. it became considerably warmer in the allerød interstadial than during the previous time intervals, despite short cold spells. common birch appeared, and gyttja and peat were formed in the shallow lakes at the beginning of the interstadial. leaching of the calcareous till increased, which resulted in lake deposition of calcareous gyttja predominantly consisting of characean algae. these algae can absorb ca2+-ions from the water and precipitate lacustrine chalk. the water level of the lakes was variable but lowered towards the end of the interstadial, again resulting in deposition of algal and detrital gyttja (fig. 20). a diversified fauna and flora existed around the lakes. during the last part of the late weichselian, younger dryas, the climate again became colder, and the water level of the lakes gradually increased. the deposits became sandier, with laminated layers of silt and clay devoid of characean algae, and with dropstones and abundant plant remains. the sedimentation documents the infill of the lakes with material from the surrounding treeless landscape. 11 700 bp the late weichselian was succeeded by the holocene with a generally warmer climate. during the different phases of the baltic ice lake from the weichselian to holocene, the sea level around møn, fakse bugt and hjelm bugt underwent marked changes, resulting in a range of different laguna-, fig. 21 the late weichselian freshwater deposits high up in the section at madses klint, which contains a bed of 20 cm thick calcareous gyttja. photo: tove stockmarr. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 25 of 53 www.geusbul let in.org coastal, fluviatile, lacustrine and bog deposits during the late weichselian (jensen 1993; bennike & jensen 1995; jensen et al. 1997). holocene deposits (postglacial, mis 1) the holocene (postglacial) after 11 700 bp is divided into preboreal, boreal, atlantic, subboreal and subatlantic time (the present). these units have been determined in a few localities, where the profiles commonly contain both freshwater and marine deposits. the most complete profile is in the borre sømose depression (fig. 22; mikkelsen 1949). part of the holocene succession has also been investigated in connection with study of the late glacial tøvelde profiles (noe-nygaard & heiberg 2001). freshwater deposits there are five different types of holocene freshwater deposits: peat (ft), gyttja (fp), clay (fl), sand (fs) and heterolithic deposits (fv). the freshwater deposits comprise c. 7% of the land area. they consist of two main age groups: those intercalated with and on top of the marine postglacial deposits and those in the isolated depressions (bogs and lakes) in the hummocky till landscape, commonly resting on late glacial deposits. figure 22 illustrates an almost continuous de positional sequence in borre sømose, from preboreal to subatlantic time, where the postglacial sequence is seen to rest on the late glacial deposits. a large number of isolated occurrences of peat are shown on the geological map. they were formed in the remnants from isolated bodies of buried ice left behind from the melting glacier. other scattered freshwater deposits were formed in shallow depressions of the flat and hummocky moraine landscapes. larger lakes such as røddinge, råby and råbylille lakes were overgrown in the holocene, and peat was formed, for example in busemarke mose. the largest of the freshwater basins is borre sømose (see the section on the holocene evolution). overlying late glacial deposits in bogs on høje møn, there are also holocene deposits, the oldest of which are of boreal age (hintze 1926, jessen in hintze 1937). as already mentioned, the freshwater deposits are dominated by peat with evidence of a preboreal, boreal, atlantic, subboreal and subatlantic vegetational development similar to that in other parts of central and southern denmark (noe-nygaard et al. 2017). large parts of a bear skeleton, most likely from the early holocene (jessen 1929), have been found in kam mose. marine deposits the marine holocene deposits form about 10% of the land area and comprise gyttja and organic mud (hp), clay (hl), sand (hs) and gravel and stones in beach ridges (hg). marine deposits dominated by beach ridges are common along the coasts. the beach ridges outline angular forelands and crescentic spits at protruding coasts and points, for example at hårbøllebro. ulvshale is a large plain of coarse-grained beach ridges underlain by marine sand deposits about 7.5 m thick with postglacial shells and a high organic content. nyord was probably formed as an overgrown foreland protected by an adjacent, 20 m high hill consisting of till (fig. 23). this hill probably consists of glacial-tectonically imbricated sheets of clayey till. flint derived from the chalk cliffs is predominant on many coarse-grained beach headlands and beach ridges. several generations of beach ridges can be found at the foot of møns klint (fig. 24) and are also known on the north coast and facing gyldenløves flak (andersen 1936). during the atlantic transgression, thick marine deposits were formed in the littorina sea. they have mostly been found in borre sømose as cardium gyttja (after the mussel cardium edule; mikkelsen 1949). the raised deposits from the littorina 0 –5 –10 2 0 100 200 300 400 500 600 m 1 3 4 5 6 7 8 9 1110 12 13 14 15 16 17 18 19 20 21 west east d ep th (m ) subboreal–subatlantic alternating gyttje layers atlantic cardiumgyttje preboreal ? claysubatlantic peat boreal–atlantic gyttje & peat weichselian tilly clay fig. 22 section through borre sømose showing the development from the late late weichselian towards the last part of the holocene. the numbers 1–21 show the positions of shallow boreholes. from gravesen et al. (2017), modified from mikkelsen (1949). http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 26 of 53 www.geusbul let in.org sea form large, almost horizontal plains at or 1 m below the present sea level, which are protected from the sea by dykes, and cultivated at kostervig. aeolian dune formation present beach ridges in the coastal areas, typically covered by aeolian dunes, can be seen on the coast at klintholm havn and råbylille as well as along the north-eastern coast of ulvshale. aeolian deposits on freshwater sediments occur in a few places. holocene development in the preboreal, in the first part of the holocene, the temperature was rapidly increasing. this was followed by a short cold spell before the climate again became warmer and dry, with migration of grasses and herbs. in the shallow lake at tøvelde, algal gyttja with a high organic content was deposited first, followed by characean gyttja when the water level increased (mikkelsen 1949; noe-nygaard & heiberg 2001). during the cold spell, the gyttja was covered by washed-out clay. when the climate again became warmer, new gyttja was deposited, and birch, pine, aspen and later beech, oak, elm and hazel migrated in. at the transition to the holocene, a shallow lake existed in the borre sø depression, which was originally a u-shaped meltwater channel from the retreating young baltic ice (fig. 22). in preboreal and boreal time, this became a bog with tree growth. in the preboreal and in the early part of the boreal, the so-called terrestrial period, a land bridge existed that connected denmark to southern sweden and southern england. the baltic ice lake still existed at the onset of the preboreal but was drained 10 300 bp. the tøvelde lakes were overgrown with plants, or tree trunks had collapsed into them during the boreal. the lakes gradually changed into bogs, and thick layers of bog peat with numerous shells of freshwater molluscs and bivalves were formed. the baltic region was again covered by a large lake, namely the ancylus lake, which was reminiscent of the baltic ice lake that had existed a few thousands of years before. in the atlantic time, large parts of denmark were inundated by the littorina sea. low-lying areas along the coasts of møn and old fjords like borre sømose became part of this sea (fig. 22; krog 1979). deposition of thick marine layers of cardium gyttja continued from atlantic into subboreal time. an extensive flora and fauna are known from this period. in subatlantic time, lacustrine and marine conditions alternated in the low-lying areas and were accompanied by alternating deposition of freshwater and calcareous gyttja. the changing depositional environment was also influenced by temporary barriers in front of fjords such as borre sømose, which in turn reflected the varying sea level. many lakes were formed towards the end of this period. these lakes grew, peat was formed, and eventually the lakes became bogs. the recent marine deposits mostly consist of sand, gravel and stones, with the latter grain sizes predominating in the beach ridges. such deposits on and below the sea floor have been explored for commercial use in hjelm bugt and fakse bugt (fredningsstyrelsen 1977, 1986). the results were compiled by jensen (1993). the funnel-shaped mouth of grønsund in the strait between falster and møn contains a central channel bordered by flats and barriers formed by storm currents. the youngest deposits are aeolian sand dunes, which are mainly located along the coasts. distribution and structure of the lithological units the position of the cretaceous chalk, the distribution of chalk sheets overlain by quaternary units and the distribution and structure of the quaternary formations fig. 23 holocene plain of beach ridges at nyord–ulfshale. view towards east. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 27 of 53 www.geusbul let in.org are documented in four cross-sections a–b, c–d, e–f and g–h through møn, which are shown on the map sheet. two additional profiles, namely a historic profile of møns klint and one of hvideklint, are included in the present volume. the four profiles are based on data from drill holes within the mapped area extracted from geus’ jupiter database (gravesen & fredericia 1984), information from ground water exploration (rambøll 2006, 2007), gravel pits (region sjælland 2012) and the coastal profiles. in this volume, the western part of cross-section g–h is reproduced as fig. 25 with the new formal stratigraphical units indicated. the profiles of the coastal exposures also contain structural interpretations of the sequences at depth. the four borehole-based cross-sections represent an overview of møn’s interior three-dimensional structure from top-chalk to surface (fig. 26). møns klint is seen in fig. 27, which is a modified version of puggard’s (1851) classic cross-section. the cross-section along the eastern 800 m of hvideklint to hjelm nakke in fig. 28 is based on a new survey supported by that of stockmarr (1996). a b fig. 24 generations of holocene beach ridges at gråryg close to møns klint. a: overview of the locality with recent beach deposits. b: details of the oldest beach ridges. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 28 of 53 www.geusbul let in.org geological cross-sections based on drilling cross-section a–b the nw–se-oriented cross-section a–b transects western møn. it runs from hvideklint in the north-west, continues through nørre frenderup, kostervig, the koster peninsula and ulvsund, and ends at gammel kalve have on sjælland. the hummocky moraine landscape at gammel kalvehave and koster land is surrounded by a holocene marine foreland. at damsholte, a sw–ne-trending valley is crossed, which contains meltwater clay, sand and gravel deposited during the ne ice advance and is interpreted as a tunnel valley (rasmussen 1965; smed 2014). this valley can be followed towards north-east through stege nor, but as is evident from the crosssection, it has been modified by the young baltic ice advance. stege nor and the hills to its south-east form a composite of parallel valleys and hills formed by the advance from the south-east of the young baltic ice. in the south-eastern part of the cross-section, there are redeformed thrust sheets of chalk in hvideklint. the chalk surface is unaffected by glacial tectonics and is approximately horizontal with an elevation of 20–30 m below sea level (fig. 2), but in cross-section a–b only at æbelnæs south-east of kostervig. cross-section c–d the n–s-oriented cross-section c–d begins in hjelm bugt at råbylille sø, continues northwards over eastern møn and over the high ground at elmelunde and reaches the north coast at nordfelt west of borre sømose. the hummocky moraine landscape belongs to the lolland till formation from the young baltic ice advance. meltwater sand and gravel of the ny borre formation are exposed around elmelunde, which has been uplifted by glaciotectonic sheets of chalk. the two above-mentioned formations can be followed all the way through the profile. in the south, holocene freshwater deposits are found in the e–w-oriented lake of råbylille sø, which is located in a depression that may be interpreted as part of the råbylille tunnel valley. the lake is overgrown and contains peat. along the coast there is aeolian sand. the almost flat chalk surface is 30–40 m below sea level throughout the profile. cross-section e–f the sw–ne-oriented cross-section e–f begins at stege and continues over damsholte, røddinge sø and store damme to fanefjord. the till landscape of stege consists of smallish elongate hills oriented sw–ne, which rest on tectonic slices of chalk. the whole profile follows a valley structure, which has drained the melting ne ice. part of the valley is interpreted as tunnel valley with small eskers including fanefjord ås and hat-shaped hills (rasmussen 1965; smed 2014). the entire valley system including the røddinge sø channel subsequently served as a drainage channel for the young baltic ice, and the ny borre formation was deposited here. the røddinge depression is now overgrown and contains peat. fanefjord is surrounded by holocene marine sand, and coarse-grained beach ridges are found along the coast of the baltic sea. sheets of chalk were exposed in the summer of 2015 when electric cables were dug in next to the transformer station outside eastern neble (fig. 26). these exposures suggest that the chalk sheets in the eastern third of the profile were probably underestimated when the profile was constructed ten years earlier. as seen in fig. 26, the chalk reaches the land surface. west of neble, the chalk surface is almost unaffected by glacial tectonics, and the pre-quaternary surface is subhorizontal at a depth of 25–30 m below sea level. cross-section g–h the longest cross-section g–h is oriented e–w through central møn. from stege it continues over råbylille, ny borre fmny borre fm lolland till fm ristinge klint till fm kraneled fmmid danish till fm mid danish till fm kobbelgård & klintholm fm lolland till fm ny borre fm klintholm till fm lolland till fm mid danish till fm ristinge klint till fm mid danish till fm ny borre fm lolland till fm kobbelgård fm chalk fig. 25 geological profile g–h, annotated with formation names and other geological units (compare with the corresponding profiles and legend on the map sheet). this and the other profiles shown on the map sheet have been constructed and interpreted based on well data from the map sheet area. meltwater sand and gravel (red), meltwater clay and silt (orange) and clayey tills (brown) dominate the quaternary deposits. the profile is orientated e–w through the central part of møn from stege to høje møn. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 29 of 53 www.geusbul let in.org råbymagle, borre sømose, busemarke, store klinteskov to møns klint at store klint and dronningestolen. part of this profile without høje møn is shown in fig. 25 with retained and new formation names indicated. in the eastern part of the høje møn area, the large chalk sheets have been pushed up to 125 m a.s.l., with a thin cover of glacial, interglacial, late glacial and postglacial deposits. adjacent to this elevated area is a marked depression with borre sømose, which drained the young baltic ice and contains lateglacial to postglacial deposits as well as sand and gravel of the ny borre formation. the area south of elmelunde is a relatively elevated hummocky till landscape. holocene marine sand was deposited along the margin of stege nor. part of the profile in fig. 25 illustrates the interpreted structure of the lithostratigraphic units in the interior of møn, which, however, cannot be directly compared with the coastal profiles because of the different nature of the underlying data. the profile shows the following formations: ristinge till, kraneled, klintholm till, kobbelgård, mid danish till, ny borre and lolland till. west of høje møn the chalk surface is subhorizontal with elevations mostly between 30–40 m and locally 20 m below sea level. it is presumed that the autochthonous chalk surface east of the chalk sheets of høje møn is located more than 100 m below sea level. isolated pockets of chalk of variable size occur in the quaternary deposits (see also houmark-nielsen 2003). chalk profiles in coastal cliffs møns klint møns klint is an impressive coastal profile, which provides a n–s-trending erosional section through the unusually steep, hilly and elevated landscape of høje møn. it has been known for over 150 years that the morphology of høje møn is due to stacked sheets of chalk (puggaard 1851; schou 1949; pedersen 2000), and the overall structure has been interpreted as a result of glacial tectonics for more than 100 years. in the last, weichselian glacial period, an ice stream advanced from the scandinavian ice cap through the baltic sea towards kattegat, and sheets of chalk and their overlying deposits were pushed up to form glaciotectonic complexes on møn and rügen in germany ( johnstrup 1874; gripp 1948). møns klint constitutes an fig. 26 elevated chalk sheets exposed in an excavation at the transformer station in neble. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 30 of 53 www.geusbul let in.org fi g. 2 7 th e m øn s kl in t p ro fil e as m ea su re d by c . p ug ga ar d at a ro un d 18 50 . http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 31 of 53 www.geusbul let in.org international key locality for glacial tectonics (fig. 27; slater 1927; aber et al. 1989; pedersen 2000, 2014). møns klint is structurally divided into a proximal imbricated complex in its southern part, an antiform stack in its central part and shallowly dipping foreland thrusts in its distal northern part (pedersen 2000). the northern part of møns klint has been especially affected by superimposed deformation. this was caused by a terminal ice advance from skåne prior to 16 000 bp when denmark became ice-free. the imbricated complex in the proximal, southern part of the cliff is characterised by a few steep sheets such as hvidskud, st. stejlbjerg and sommerspir. these structures are shown in the block diagram on the geological map (the drawing ’stratigraphy and structures of møns klint’). note that the elevation of the décollement surface is c. 100 m below sea level (fig. 42; pedersen 2000). the southernmost part of the complex is jættebrinken, which consists of three imbricated chalk sheets that are now flat-lying (fig. 27). glacial sediments occur in a depression south of jættebrinken, where the chalk surface is below sea level. a path from the lighthouse to the sea at the southern point of møns klint traverses the glacial deposits, which are however concealed by dense shrubbery. the southern part of møns klint is interpreted as the location where the advancing ice was in direct contact with the hinterland of the imbricate thrust sheets. the till deposited here is interpreted as belonging to the lolland till formation. it acquired an increased thickness during prolonged deposition in the tectonic depression behind the glaciotectonic complex. hundefangsfaldet separates jættebrinken from hundefangsklinten farther north, which is the first of the steep cliffs in the proximal zone (figs 27, 42). this zone continues to sommerspirpynten, where the wide ‘gutter’ of sandfaldet has acquired its name from abundant local meltwater sand. sandfaldet is interpreted as a piggyback basin, where a river of meltwater flowed through the valley between the sommerspir sheet and the next sheet north of sandfaldet, while the sheets were being moved forward by the ice pressure. the sheet north of sandfaldet extends from græderen to maglevandsfaldet (fig. 27). along its flanks, the sand overlies the lower side of the sommerspir sheet and the upper side of the sheet at græderen. the sand is mildly folded into a syncline as a drag structure along the imbrication thrust. these features show that the sand was deformed by contemporaneous glaciotectonic processes more or less contemporaneously with its deposition. a detailed structural study of the area below geocenter møns klint (pedersen 2003; pedersen & grave sen 2009) has shown that the structures at dronningestolen resulted from overthrusting of several flat-lying sheets in the early phase of the deformation (figs 27, 42). this was followed by collective transport over a more deepseated ramp, whereby the sheets were folded into a hanging-block anticlinal structure (pedersen 2000). the thrust sheets in the distal part of the glaciotectonic complex are moderately inclined towards the south (fig. 27, 42). field work by the authors has shown that they extend into the foreland, which consists of kobbelgård formation clay. an unusual roll-over anticline is exposed below nonnebrinken. tectonic brecciation thrust fault lolland till formation mid danish till formation møns klint formation klintholm till formation kraneled formation 0 20 m –20 –40 –60 –80 hk 11 hk 13 hk 12 hk 10 hk 9 hk 8 hk 7 hk 6 hk 5 hk4 hk3 hk2 hk1 sv ne 0 20 m -20 -40 -60 -80 100 m 0 20 m -20 -40 -60 -80 hk11 hk13 hk12 hk10 hk9 hk8 hk7 hk6 hk5 hk 4 hk 3 hk 2 hk 1 sw nø 0 20 m –20 –40 –60 –80 40 80 1000 20 hk5 hk4 stubberup have formation kobbelgård formation fig. 28 the structure of hvideklint based on new measurements supplemented by information from stockmarr (1996). hk1 to hk12 refer to individual chalk sheets numbered from northeast to southwest. these are also used in figs 29–39 from hvideklint. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 32 of 53 www.geusbul let in.org hvideklint the more than 1 km long and about 20 m high coastal cliff of hvideklint faces the western part of hjelm bugt. the profile is oriented ne–sw and provides an erosional section of the hummocky landscape of south-western møn (fig. 28). the cliff, which appears as a shiny white stretch from a distance, consists of more than 13 chalk sheets with intercalated glacial deposits. the overall structure of hvideklint and its interior architecture are controlled by three glaciodynamic factors: (1) sheets of chalk were first pushed up from south-east towards north-west by the ne ice advance. (2) the thrust sheet complex was then deformed by the young baltic ice advance from south-east towards northwest. (3) the deformation of the chalk-clay association resulted in extensive brecciation and formation of mud diapirs (aber 1979; berthelsen 1979; stockmarr 1996). the cross-section of the hvideklint structures shown in fig. 28 is based on a survey of the cliff, strike and dip measurements of depositional layers and thrust sheets, measurement and construction of fold axes and volumetric considerations (balanced crosssections, pedersen 2005). the structural interpretation is supported by the constructed top of chalk map of fig. 2 and the cross-section e–f on the map sheet, both of which are based on borehole data mostly from western møn. a b fig. 29 folds at hvideklint. a: overturned syncline in sheet hk3 with an e–w-oriented fold axis. the light glaciofluvial sand of the stubberup have formation is folded into the darker clay of the kobbelgård formation. b: the nose of the overturned syncline includes alternating layers of clay, silt and sand. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 33 of 53 www.geusbul let in.org the cross-section also acts as a model for the hummocky landscape around hjelm and nørre frenderup. the cross-section begins at a small stream crossing the beach below ørbæklund. the north-easternmost c. 300 m of hvideklint (hk) are dominated by the chalk sheets hk1 and hk2. a narrow, overturned syncline on top of hk1 contains the lower units of the kobbelgård formation. the folding has also affected the uppermost part of the underlying chalk, which has been converted to a chalk-glacitectonite. the folded material is truncated by a thrust overlain by the mid danish till formation. the lolland till formation discordantly overlies the entire coastal profile. the thrust contact itself between hk1 and hk2 is not always well exposed but can nevertheless be identified where the south-western flank of the up-thrust hanging-wall block anticline (hk1) anticline is almost perpendicular to the top of the shallowly ne-dipping layering in the upper part of hk2. increasing glaciotectonic brecciation obscures the larger structures towards the nose of hk2, which has been pushed up along a c. 45° ne-dipping thrust, where the clay of the kobbelgård formation has been inverted along the south-western flank of the anticline. a 120 m stretch of slumped and overgrown glacial deposits separates the chalk exposure at the nose of hk2 and another exposure showing the top of ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ hk7 chalk chalk thrust fault kraneled fm hk6 erosion surface fig. 30 thrust fault at hvideklint separating sheets hk6 and hk7 with an isoclinally folded syncline consisting of kraneled formation clay. the thrust fault can be traced throughout the photo. hk8 hk7 hk8 hk9 fig. 31 thrust faults between sheets hk7, hk8 and hk9 at hvideklint together with the front of sheet hk9. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 34 of 53 www.geusbul let in.org hk3. an overturned syncline in the north-eastern part of this stretch has an e–w-oriented fold axis, which indicates thrusting from the south (fig. 29). the syncline and the accompanying thrust dip steeply to the south. accordingly, the fold structures related to the ne-directed thrusting from north-east have been succeeded and redeformed by n-directed thrusting from the advancing young baltic ice. the syncline overturned towards the north is clearly visible in the laminated and thinly bedded meltwater sediments of the kobbelgård formation (fig. 29). in the same stretch towards the top of hk3, there are three to four imbricated thrust sheets dominated by the clay of the kraneled, klintholm till and kobbelgård formations. since there are no exposures of chalk between these sheets, they are interpreted as having been carried tectonically piggy-back on hk3. it is es timated from the interpreted profile that these sheets are rooted in the top of chalk unconformity, 35 m below sea level. considering that the maximal thickness of the exposed sheets is about 15 m, it is expected that the detachment zone of the chalk slices themselves is at least 15 m deeper. it is indicated in the profile that the detachment zone is located 80 m below sea level, but at this depth only chalk would be thrust over in situ chalk. hk8 fig. 32 intensely deformed chalk in thrust sheet hk8 at hvideklint, grading from chalk-glacitectonite to brecciated chalk-till deposits. hk10 hk9 ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ fig. 33 folded and thrust-displaced layers at hvideklint, representing parts of the kraneled, klintholm till and kobbelgård formations between the sheets hk9 and hk10. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 35 of 53 www.geusbul let in.org the chalk slices hk3, hk4 and hk5 form a contiguous complex, which may be interpreted as an antiformal stack. that is to say, the slices were first pushed laterally on top of each other and then thrust over an underlying ramp below hk5. a small separate chalk wedge transects the nose of hk4 and is interpreted as having been pushed up from the south. hk6 is a thin slice of chalk, which has been thrust over the relatively thick chalk sheet hk7. the separating thrust is associated with an isoclinally folded syncline of dark clay, which presumably belongs to the kraneled formation (fig. 30). the chalk of hk6 and the upper part of hk7 is broken up and brecciated and is best described as a chalk glacitectonite. the nose of hk7 has been thrust over the top of hk8, where several quaternary layers have been folded into an overturned syncline under the base of hk7. hk8 is a relatively short sheet, carried on the back of chalk sheet hk9. a still smaller and thinner sheet of chalk glacitectonite has been dragged discordantly across the noses of hk8 and hk9 at the base of the mid danish till formation. the tip of hk9 is so intensely affected by shear deformation that is intermediate between a chalk glacitectonite and a brecciated chalk till (figs 31, 32). between the chalk sheets hk9 and hk10, there are two sheets of disrupted and folded units from the kraneled, hk10 lolland till fm mid danish till fm klintholm till fm fig. 34 cretaceous chalk at the front of sheet hk10 at hvideklint, overlain by a succession of quaternary deposits including the kraneled, klintholm till, kobbelgård, mid danish till and lolland till formations. mid danish till fm hk 12 hk 11 kobbelgård fm lolland till fm fig. 35 the steeply dipping chalk sheets hk11 and hk12 at hvideklint, truncated by the almost horizontally deposited lolland till formation including a lens of chalk-glacitectonite. person for scale in centre of photo. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 36 of 53 www.geusbul let in.org klintholm till and kobbelgård formations. the upper of these sheets has been folded into an overturned anticline, and the layers along their mutual thrust contact have been intensely deformed by reverse faulting (fig. 33). dark clay occurs on the top of hk10, which can be assigned to the kraneled and klintholm till formations. strongly reversely displaced, sandy layers belong to the kobbelgård formation. three different lodgement tills occur at the frontal part of hk10; sandy and clayey sediments of the kobbelgård formation separate the tills of the klintholm till formation from those of the mid danish till and lolland till formations (fig. 34). the chalk sheets hk11 and hk12 are relatively thin and have steep orientations (fig. 35). because of shear deformation there are both layers and lenses of black clay in chalk and layers and lenses of chalk in dark clay (figs 36, 37). towards south-west from hk11, there are increasing occurrences of clay, mainly of the kraneled and kobbelgård formations, which have caused in creased sliding and blurring of the coast profile (fig. 38). an additional chalk slice, hk13, has been identified farther south-west towards hjelm nakke. however, the tectonic structures are so complex that is meaningless to construct a profile. an example of the complex de for mation patterns can be fig. 36 thin layers of chalk-glacitectonite at hvideklint, which were squeezed up above the existing pile of chalk sheets hk1–hk12. the thin chalk layers and lenses have been sheared into the dark clay of the kobbelgård formation. fig. 37 lenses and thin layers of clay in the chalk of sheet hk8 at hvideklint, deformed by thrust faulting. the chalk has completely lost its primary sedimentary layering and has been transformed into a glacitectonite. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 37 of 53 www.geusbul let in.org observed in mud diapirs, which occur where water-saturated clay deposits have been subject to glaciotectonic deformation (fig. 39). de for mation patterns can be observed in mud diapirs, which occur where water-saturated clay deposits have been subject to glaciotectonic deformation (fig. 39). glaciodynamic sequence stratigraphy pedersen (2012a) proposed a glaciodynamic sequence stratigraphy for terrestrial glacial deposits, which both includes the formation of glaciotectonic complexes and the deformed lithologies themselves (the so-called glacitectonites, pedersen 1988, 2014; pedersen et al. 2018). a glaciodynamic sequence corresponds to a glaci odynamic event. the base of a glaciodynamic sequence is the deepest level affected by the corresponding event. in a complete glaciodynamic sequence, this is the basal detachment level of the glaciotectonic deformation, i.e. the most deep-seated level where glaciotectonic offset occurs. the individual units within the complex are found above this level. they may consist of pre-quaternary units such as glacially deformed chalk, as well as pre and proglacial formations such as meltwater deposits. fig. 38 brecciated clasts of chalk in mud along the sole of the thrust faults at hvideklint. the large mud content in the kraneled and kobbelgård formations commonly blurs exposures of the south-western part of the cliff. fig. 39 mud diapirism at hvideklint formed during glacial tectonic deformation of water-saturated clay deposits. the exposure is almost 2 m high. photo: tove stockmarr (1996). http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 38 of 53 www.geusbul let in.org the base of such a complex is an angular discordance, above which a glacitectonite is formed at the base of a lodgement till. the top of the glaciodynamic sequence is the top of the lodgement till that was deposited during the ice advance that caused the glaciodynamic event (fig. 40). four different glaciodynamic sequences can be differentiated on møn, namely klintholm, hvideklint, møns klint and aborrebjerg. klintholm glaciodynamic sequence the coast profiles east and west of klintholm contain dislocated units of meltwater clay and sand as well as diamict deposits, which may either be lodgement tills or dropstone tills (see discussion on the formation of diamict deposits in houmark-nielsen 1994). variable amounts of chalk and chalk sheets have been worked into the basal parts of the glacial deposits. the chalk sheets have been pushed up from east to west and are discordantly overlain by dark grey clayey till. this sequence is named the klintholm glaciodynamic sequence. the corresponding glaciodynamic event took place during the earliest (ristinge) ice advance in the early to middle weichselian c. 55 000–50 000 bp. in terms of climate, the event was followed by the sandnæs mild period. hvideklint glaciodynamic sequence the glaciotectonic complex of hvideklint consists of about 15 sheets of chalk above the top of chalk unconformity and includes variable amounts of glacial material attached to the upper surfaces of the sheets. all the sheets are truncated by the overlying mid danish till formation, commonly with a smeared-out unit of chalk glacitectonite 0.5–1 m thick at its base. this sequence is named the hvideklint glaciodynamic sequence. it was formed by the predominant ne ice advance, which spread out from a source area in central sweden to the main stationary line in denmark, northern germany and poland. on møn, the ice advance was directed from north-east to south-west. also, the second phase of the præstebjerg hill complex is included into this glaciodynamic sequence. the glaciodynamic event took place in the latter part of the weichselian glaciation during the lgm. møns klint glaciodynamic sequence the glaciotectonic complex of høje møn is a contiguous glaciotectonic unit that was pushed up from south towards north. its base is a detachment surface c. 100 m below sea level (pedersen 2000, 2003; pedersen & gravesen 2009); its depth was estimated using the principles of balanced glaciotectonic cross-sections described by pedersen (2005, 2006). above this surface, the chalk is no longer in situ. the glaciotectonic complex has a classic internal structure comprising, from south to north, a proximal, a central and a distal part. the tops of the chalk slices and their intercalated glacial deposits are truncated by the lodgement till of the lolland till formation, commonly including a dragged base of chalk glacitectonite. this sequence is here named the møns klint glaciodynamic sequence. the associated event is ascribed to the rapidly advancing young baltic ice at about 18 000 bp, which was also responsible for phase 3 and possibly 4 of the præstebjerg hill complex (see the chapter on landscape formation). the glaciodynamic formation of møns klint has controlled the overall structure of the entire island of møn and particularly the coastal area facing hjelm bugt (fig. 41) and the superimposed deformation of hvideklint. aborrebjerg glaciodynamic sequence the northern half of the glaciotectonic complex of møns klint has been superimposed by glaciotectonic thrusting from north-east (fig. 42; pedersen 2000). the best example of this is the chalk sheet at vitmunds nakke, which has clearly been pushed up from north-east towards south-west and which also displaces the lolland till and mid danish till formations and older glacial units. an analysis of the superimposed folds shows that aborrebjerg itself constitutes the culmination of a fold of a fold (fig. 43). the sequence is therefore named the aborrebjerg glaciodynamic sequence. the thrusting corresponds to an ice advance from the area north-east of møn, where the receding young basal till glacitotectonite thrust fault sand & clay décollement chalk d c b a ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ fig. 40 the concept of glaciodynamic sequence stratigraphy (after pedersen 2012a). a: location of the basal décollement zone in the upper part of the chalk. b: formation of thrust fault imbricates consisting of chalk and glacial deposits. c: truncation of the glacier sole during formation of glacitectonite. d: final deposition of lodgement till above the glacitectonite. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 39 of 53 www.geusbul let in.org baltic ice must have extended a ‘piedmont’-shaped ice tongue during its general meltback. a sw-facing ice front displaced the already previously displaced chalk sheets. according to the established models for the meltback of the young baltic ice, this re-advance must have taken place 16  000 bp (houmark-nielsen 2010, houmarknielsen et al. 2017). landscape features and their formation geomorphological description the following geomorphological description is based on the geological map and a terrain model of møn (fig. 44). glaciotectonic landscape with composite ridges the most characteristic morphological feature of høje møn is the elongate, parallel ridges separated by long and narrow valleys. the ridges consist of glaciotectonic thrust sheets of chalk. the chalk is commonly present at the surface in store klinteskov and best observed after major storms, when the chalk becomes exposed around the roots of fallen trees. in the southern part of klinteskoven the ridge orientation is e–w. characteristic examples are gråryg with the adjacent elongate valley of nælderende on its northern flank. farther south, the ridge orientation turns towards ese–wnw, as in the valley of the siesøerne lakes and the adjacent siesø bjerg to the north. the ridge that extends into the wood from dronningestolen may also be considered part of the e–w-oriented hill system. farther north, there is a marked change in the ridge orientation to se–nw. a few of the ridges close to the cliff reach a sse–nnw orientation, which can also be observed in the lake depressions from hunesø to liselund. based on the glaciodynamic interpretation of the møns klint complex, the hills can be classified as composite ridges. their changing directions are interpreted as resulting from an initial pressure from the south-east by the young baltic ice, which created the ridges, followed by an ice advance from skåne in the north-east towards the end of the weichselian, which redeformed and to some degree also re-oriented the ridges (pedersen 2000). an e–w ridge orientation is also observed in the glacial morphological landscape around stege nor, which can itself be considered a morphological element that follows the orientation of the ridges. the highest points in the landscape are interpreted as corresponding to places where the underlying chalk was already pushed up from north-east. moraine plateaus about 70% of the surficial material on møn consists of clayey till, which almost entirely belongs to the lolland till formation. apart from store klinteskov and a few other elevated hill areas this landscape can be characterised as a lodgement till landscape forming a flat, a b fig. 41 block diagrams showing the formation of møn’s hilly landscape in front of the young baltic ice. a: the young baltic ice occupying the depression in hjelm bugt. b: exposure of an arcuate, hilly landscape during meltback of the ice. note that the hills are cut by valleys formed by meltwater spillways from the adjacent ice margin along the south coast of møn. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 40 of 53 www.geusbul let in.org a b f1 aborrebjerg was created by two superimposed hanging-wall anticlines vitmunds nakke formed by the steeply ne-plunging f1-fold f1 f1 f2 fig. 43 formation of aborrebjerg and vitmunds nakke. aborre bjerg is the highest point of møn and was formed during two separate events: a: superimposed thrust fault deformation by the young baltic ice advance from the south (f1). b: superimposed faulting by the latest ice advance from southern sweden (f2). imbricate fan ramps antiform stack slotsgavle foreland near thrusting maglevandsfald ramp dronningestolen antiformal stack gråryg thrusting thrusting from enejydelejet superimposed thrusting young baltic ice advance from sse fig. 42 the structural evolution of møns klint (after pedersen 2000). blue: chalk. brown and orange: glacial deposits. red: thrust faults. blue arrows show the directions of ice push from the sse and ene, respectively. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 41 of 53 www.geusbul let in.org levelled terrain with scattered lows without outlets. the lows are depressions left by former lumps and sheets of ice buried in the clayey till during the young baltic or earlier ice advances (fig. 45). the till is underlain by meltwater sand of the ny borre formation. the sand appears in the valleys, as easily observed east and west of stege nor. towards the south-west from stege nor, the sand is widespread in the depression around damsholte and røddinge sø and in the valley extending towards fanefjord. hat-shaped hills and superimposed eskers on western møn, several hills may be interpreted as hatshaped hills or eskers, which have been affected by later ice advances than those responsible for their original proand intraglacial deposition. the next sections describe the processes that formed the proglacial deposits and their subsequent deformation along the ice margin. the intraglacial deposits were formed in tunnel valleys under the ice or in large ice-dammed lakes with margins and bases associated with fracture zones in the ice. after the meltback, such deposits stand out in the landscape as hills of sand and gravel, which are described in the older literature as hat-shaped hills (berthelsen 1980). it remains uncertain how many of the hills in western møn can be considered hat-shaped hills, but one of them, præstebjerg, has been exploited as a gravel pit, which allowed a study of its interior structure during its excavation (fig. 46; berthelsen 1980). the main criterion that defines a hat-shaped hill is that it stands out as an isolated hat in the landscape. it is also required that the internal layering is deformed, oblique and steeply dipping. this criterion is ascribed to berthelsen (1980), who excluded a large number of hills with this exterior shape but without dislocations, for example on langeland. an investigation of præstebjerg by berthelsen (1980) demonstrates that the depositional and deformational processes in hat-shaped hills may be very complicated. in the first phase of his interpretation, the material in the præstebjerg hill was deposited as a proglacial meltwater plain, which was then folded by an ice advance from south-east, probably the klintholm ice advance (fig. 47); without a detailed osl or infrared stimulated (irsl) age determination it is impossible to a b fig. 44 geomorphological elevation model of møn. a: viewed towards north. b: viewed towards west. yellow-green: 0–25 m a.s.l. brown: 25–50 m a.s.l. white: above 50 m a.s.l. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 42 of 53 www.geusbul let in.org establish if the ice advance represented the ice of the ristinge klint till formation or that of the klintholm till formation, although the layers have clearly been deformed towards north-west by a baltic ice. the second phase (phase 2 in fig. 47) is that of the overriding ne ice. just before the lodgement till of the mid danish till formation was deposited, the sandy hill was deformed by a sw-directed ice pressure. the third phase began with strong erosion of the northern hill flank by a meltwater stream, which subsequently deposited a thick package of meltwater sand and gravel. the deposition took place in front of an ice advance from the south-east, which then covered the hill and deposited a thin layer of lodgement till. in the fourth and final phase, the hill was covered by a second ice advance from south-east, which folded the layers inside the hill on a ne–sw-trending fold axis and overturned the folds towards nw (fig. 47). it is suggested in this interpretation that the two latest ice advances were oscillating and related to the ice stream that deposited the lolland till formation. marginal moraine landscape and meltwater plains a large accumulation of meltwater sand and gravel occurs from ny borre southward to busemarke. two of møn’s largest gravel pits were also found in this area at ny borre and in the western outskirts of busemarke. towards the south, the area is bounded by the marginal moraine hills north and south of råby sø and busemarke mose, with the most prominent hilltop located at kobbelgård. together with other hill stretches parallel to the coast of hjelm bugt, these hills form a terminal moraine landscape. in front of the hills, a meltwater plain stretched from busemarke to ny borre. farther north, erosion by the same meltwater created the broad depression of borre sømose. the american ex pression for a landscape formed by melt water is ‘spillway’, which may be applied to the valley towards maglemose. the depression stretching from the north coast to borre sømose as well as the maglemose depression were also spillways for the meltwater when the retreating ice margin experienced a short-lived standstill at the coast of hjelm bugt. at this time, the meltwater mainly flowed through the valley towards fig. 45 the boundary between the elevated landscape of høje møn and the landscape towards the west of the lower, undulating till plain. the white cliffs of hvideklint can just be seen in the horizon on the far side of hjelm bugt. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 43 of 53 www.geusbul let in.org a b fig. 46 the hat-formed hill of præstebjerg on western møn. a: the hill viewed from the south, with excavated material at its top. b: vertical layers of meltwater sand and gravel inside the hill, cut and overlain by clayey till of the lolland till formation. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 44 of 53 www.geusbul let in.org maglemose, where sand and gravel were redeposited during the late glacial meltback. similar processes also controlled the areas where stege nor enters stege bugt as well as at klostervig, røddinge sø and fanefjord. in total, meltwater sand accounts for nearly 10% of møn’s surficial deposits. when the young baltic ice finally melted away from møn, the island was a terminal moraine landscape. hjelm bugt forms a lobe-shaped basin, which may be considered the source area for møn’s quaternary deposits. marine foreland with beach ridges and spits around 8000 bp, the inner danish waters were inundated by the atlantic transgression, and sea level increased by more than 20 m. the transgressive marine deposits are located in low-lying coastal areas and form c. 10% of the deposits on møn. it was at this time that the bogs of borre sømose and maglemose were transformed into fjords, and klostervig was a narrow strait between small islands and islets. these fjords and straits were gradually sealed off by beach ridges fed by sediment transport along the coast. the lagoons behind the ridges were overgrown and changed into bogs with thick layers of peat. the development of the marine foreland became most prominent where it was built out by sediments transported by marine currents. ulvshale was created in this way. mapping of beach ridges developed into crescentic bars clearly indicate a westward transport direction along the north coast of møn. at the northern point of the island, the marine foreland was built out towards north-west, where the beach ridges are prominent and commonly superimposed by smaller dunes. a smaller area at hårbølle havn is also a new marine foreland. the patterns of the beach ridges show that an angular foreland was formed by sediment transport in two opposite directions, namely southward through kraneled & klintholm till fm mid danish till fm older glacio�uvial sand ny borre fm lolland till fm 4 3 2 1 fig. 47 model for the formation of præstebjerg. 1: folding of the kraneled and klintholm till formations during the klintholm ice advance from eastsouth-east. 2: superimposed folding and deposition of the mid danish till formation by the ne ice advance from northeast. 3: glaciofluvial erosion by meltwater streams moving from east to west and deposition of the ny borre formation. 4: superimposed deformation due to the young baltic ice advance from south-east and deposition of the lolland till formation. partly from berthelsen (1979). http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 45 of 53 www.geusbul let in.org grøn sund and north-westward from hjelm bugt through grønsund. the transport of material to the northwestern and south-western corners of møn indicates a source area located on the eastern side of the island. it is naturally the steep cliffs of møns klint that have supplied erosional material, not least from the many rockfalls and landslides described next. rockfalls and landslides rockfalls only occur in a few places in denmark, and møns klint is the most prominent area because of its high and steep chalk cliffs (pedersen et al. 1989; nadim et al. 2008; pedersen & gravesen 2009; pedersen 2012b). landslides, which occur in many parts of denmark, are also common at møns klint (nadim et al. 2008). the difference between rockfalls and landslides is lithology and topography. rockfalls occur where a more or less consolidated lithology slides or falls from a steep or nearly vertical cliffside. landslides occur where the lithology is water-saturated clay, and a terrain gradient of just a few degrees is sufficient. the sliding body moves slowly or moderately quickly forward on a thin film of mud at high pore-water pressure. landslides range from mud streams to terrace-shaped block displacements. rockfalls in the last 100 years, only one danish rockfall has been reported to have caused a death. this occurred on 29th july 1994 from maglevandsnakke just south of dronningestolen (pedersen 1994), and surprisingly this rockfall took place on a dry summer day. the size of the fallen block was only a few thousand cubic metres (2500–3000 m3, pedersen 1994), but the risk of a fatal accident was high inasmuch as the rockfall occurred at a time when many tourists were on the beach. one person was killed, and since then this part of the cliff has been carefully monitored. however, it is difficult to predict where another rockfall might take place. a survey of rockfalls in the last 50 years suggests that a large rockfall occurs about every 5 years (table 1). the latest took place in january 2007 at taleren cliff, where 100 000 m3 of material was spread out 300 m into the sea, (fig. 48; pedersen 2007). that year, there were also several other rockfalls at møns klint and at other danish localities. the reason for this was an unusually wet autumn, where the average groundwater table was about 1 m higher than normal. accordingly, the pore-water pressure near the surface was very high, which reduced the rock strength and in turn facilitated jointing. once joints have formed, the conditions for a rockfall have been created. fig. 48 landslide at store taler in january 2007. aerial photograph. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 46 of 53 www.geusbul let in.org every time a large rockfall at møns klint has taken place, it has resulted in a peninsula of chalk, and the fresh lumps of chalk are favoured targets for fossil hunters (fig. 49). the area immediately below the broken cliff is usually sealed off, because smaller blocks with sizes of a couple of cubic metres are commonly released up to two days after the initial rockfall. a survey of the erosion rate along møns klint provides an overview of rockfall conditions. a comparison of the position of the coastline in 1899 and 100 years later (fig. 50) provides a good overview. the average annual erosion is 30 cm, which is about twice as much as other danish coastal cliffs. however, the localities with the highest erosion, 44–49 cm/year, is not below the highest chalk cliffs but at the falls between them. for example, there is a continuous tendency for sliding at sandfaldet between the sommerspir cliff and græderen south of dronningestolen, where a spring emerges, and the sand is relatively poorly compacted. landslides landslides of cliffs containing layers of clay are common along most danish coastal cliffs. on møn, they are most common in the northern part of møns klint. the largest known single event took place in 1905, where 4 ha of liselund park slid into the sea after a major storm. fifteen years later, the activity at liselund was re-activated with hundevangsfaldet fruerstuefaldet skredrenden tragten sandfaldet maglevandsfaldet maglevandsfaldet sandskredsfaldet sækkepibefaldet gukkenhule dronningestolen 24 7 33 11 33 15 5 4 21 29 29 33 32 39 46 35 24 48 6 34 41 33 24 36 36 29 29 29 33 38 34 34 40 41 35 35 36 32 44 21 21 17 coast c. 1999 coast c. 1889 200 m n s erosion (m)35 fig. 50. erosion and retreat of møns klint over time by landslides and erosion, shown as the calculated difference between the coastlines in 1899 and c. 1999. fig. 49 after rockfalls from the cliffs, fossil hunting in the displaced chalk material is very popular. photo: peter moors. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 47 of 53 www.geusbul let in.org fig. 51 landslide on the beach below liselund castle with clay from the kobbelgård formation and fallen trees. extrapolated position of shoreline before landslide elevated cli� foot after landslide elevated shoreline shoreline elevated by landslide, september 2007 fig. 52 landslide at abildgaards fald in northern part of møns klint in september 2007. the toe of the slide displaced the beach plane. the three lines mark the elevated cliff foot, the elevated former shoreline and its recent position with an extrapolation. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 48 of 53 www.geusbul let in.org a landslide of half the previous size and was resumed again in 1927. a terrace morphology and gullies from the previous landslides (hintze 1904, 1920) are still visible today. the underlying reason for all these landslides can be observed on the beach at liselund, where bluish grey ice-lake clay of the kobbelgård formation is exposed at the foot of the cliff. the landslides take place on top of this clay. fallen trees on the beach, which consists of flint-rich gravel and stones, show that landslides are still active (fig. 51). the clay of the kobbelgård formation has also caused landslides in adjacent coastal cliffs. in the autumn of 2007, a cylindric landslide took place at abildgaards fald, whereby the beach level at the foot of the slide was increased by 6 m (fig. 52). the entire coastal cliff northwards from liselund via hellehavn nakke to pomlerende is dominated by landslides, which displace the quaternary deposits that characterise the northern part of møns klint. acknowledgements we are very grateful to tove stockmarr and gerald hyde for permission to use their cand. scient. (m.sc.) theses including figures and photographs. we thank gunver krarup pedersen for many constructive scientific com ments to an earlier version of the text, and the two external reviewers, helena alexanderson and nikolaj krogh larsen, are thanked for their comprehensive reading and suggestions of improvements. the authors remain responsible for the geological map sheet and the final text. table 1 større fjeldskred og jordskred ved møns klint mellem 1801 og 2015§ year date locality size class† 1801 taleren 1868 24/12 part of dronningestolen 3 000 000 m3 3 1899 sækkepibe fald 100 000 m3 3 1904 31/12 north of sommerspiret and gukkende hule 1905 05/11 liselund muddy landslide 4 ha 3 1910 13/03 hundefangsklinten 1914 01/01 dronningestolen and forchhammers pynt 1920 14/09 liselund muddy landslide 2 ha 3 1927 05/11 large fall at liselund 1927 22/11 slotsgavlene, pomlerende, nonnebænken, stubberup have 1928 11/05 hundefangsfaldet 10-15 ha 3 1929 southern part of dronningestolen 1939 25/01 puggaards pynt temporary peninsula 2–300 m across 3 1940–46 gukkende hule, part of vidmunds nakke, forchhammers pynt 1948 15/01 dronningestolen 1952 24/01 røde udfald and parts of vitmunds nakke to puggaards klint collapsed temporary peninsula 500 m across 3 1958 vitmunds nakke 1963 maglevandsfaldet 1970 store stejlebjerg 2 1979 cliff face of sommerspiret collapsed 2 1980 hylledalsfald 1 1988 13/01 sommerspiret 1 1994 29/07 southern part of dronningestolen. french tourist killed at 12:30 3000 m3 1 1998 04/03 freuchens pynt, 100 m from the coast 3 2003 05/07 store stejlbjerg, at 8:50 12 000 m3 2 2004 16/01 northern part of jættebrinken, at 16:00 4000 m3 2 2004 17/03 sommerspirpynt 2000 m3 2 2007 27/01 store taler collapsed in the evening 100 250 m3, peninsula 300 m across 3 2007 02/03 freuchens pynt, afternoon 25 000 m3, peninsula 150 m across 3 2007 30/09 abildgårds fald, evening, muddy landslide, toe of slide and beach plane elevated 6 m 2 2009 15/10 nylands nakke, evening. clear sky during autumn high pressure gave way to night frost conditions. rockfall of existing overhang 5000 m3 1 2015 05/12 the cliff peak, græderen, collapsed under storm bodil 1 § the table includes rockfalls of cretaceous chalk as well as landslides of quaternary deposits. 1: small. 2: medium. 3: large. http://www.geusbulletin.org pedersen & gravesen 2022: geus bulletin 51. 8336. https://doi.org/10.34194/geusb.v51.8336 49 of 53 www.geusbul let in.org additional information funding funded by geus. additional files the møn geological map sheet is supplied as an additional file at https://doi.org/10.22008/fk2/tmfh5w references aaris-sørensen, k., petersen, k.s. & tauber, h. 1990: danish finds of mammoth (mammuthus primigenius (blumenbach)): stratigraphical position, dating and evidence of late pleistocene environment. danmarks geologiske undersøgelse serie b 14, 44 pp. https://doi.org/10.34194/serieb.v14.7081 aber, j. 1979: kineto-stratigraphy at hvideklint, møn, denmark and its regional significance. bulletin of the geological society of denmark 28, 81–93. aber, j.s., croot, d. & fenton, m.m. 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https://doi.org/10.1016/j.quascirev.2007.06.021 https://doi.org/10.34194/raekke4.v2.6980 descriptive text to the geological map of denmark, 1:50 000, møn 1511 i, 1511 iv and 1512 ii abstract topography methods history of geological mapping on and around møn pre-quaternary deposits late cretaceous chalk glacial and interglacial deposits data and interpretation older quaternary deposits: saalian, eemian, early and early middle weichselian (mis 6, 5e, 5a–5d and 4) saalian (mis 6) eemian (mis 5e) early and early middle weichselian (mis 5a-5d, mis 4) middle and late weichselian (mis 3, 2) ristinge klint till formation (mis 3) kraneled formation (mis 3) klintholm till formation (mis 3) kobbelgård formation (mis 3) stubberup have formation (mis 3) mid danish till formation (mis 2) ny borre formation (mis 2) lolland till formation (mis 2) late weichselian (late glacial, mis 2) late weichselian development (late glacial) holocene deposits (postglacial, mis 1) freshwater deposits marine deposits aeolian dune formation holocene development distribution and structure of the lithological units geological cross-sections based on drilling cross-section a–b cross-section c–d cross-section e–f cross-section g–h chalk profiles in coastal cliffs møns klint hvideklint glaciodynamic sequence stratigraphy klintholm glaciodynamic sequence hvideklint glaciodynamic sequence møns klint glaciodynamic sequence aborrebjerg glaciodynamic sequence landscape features and their formation geomorphological description glaciotectonic landscape with composite ridges moraine plateaus hat-shaped hills and superimposed eskers marginal moraine landscape and meltwater plains marine foreland with beach ridges and spits rockfalls and landslides rockfalls landslides acknowledgements additional information references figures fig 1: index map showing position of the map sheet møn. see caption for full description. fig 2: top of chalk below the quaternary deposits. see caption for full description. fig 3: photo of pale grey chalk with black flint layers. see caption for full description. fig 4: palaeogeographic map of the late cretaceous sea in northern europe. see caption for full desc fig 5: stratigraphic divisions. see caption for full description. fig 6 a: schematic stratigraphic log and a locality map. see caption for full description. fig 6 b: schematic stratigraphic log and a locality map. see caption for full description. fig 7 a: eemian deposits at hjelm nakke. see caption for full description. fig 7 b: eemian deposits at hjelm nakke. see caption for full description. fig 8: composite section showing the thicknesses and horizontal distribution of the lithostratigraph fig 9: photo of the ristinge klint till formation at store stejlebjerg, møns klint. see caption for fig 10: photo of the ristinge klint till formation. see caption for full description. fig 11: photo of the klintholm till formation. see caption for full description. fig 12: photo of the kobbelgård formation. see caption for full description. fig 13: photo of the kobbelgård formation. see caption for full description. fig 14 a: photo and sedimentological log of the stubberup have formation. see caption for full descr fig 14 b: photo and sedimentological log of the stubberup have formation. see caption for full descr fig 15: photo of the mid danish till formation. see caption for full description. fig 16: sedimentological log of the ny borre formation. see caption for full description. fig 17: photo of the ny borre formation. see caption for full description. fig 18: photo of the lolland till formation. see caption for full description. fig 19: photo of the lolland till formation. see caption for full description. fig 20: photo of late weichselian – holocene deposits. see caption for full description. fig 21: photo of late weichselian freshwater deposits. see caption for full description. fig 22: section through borre sømose. see caption for full description. fig 23: photo of holocene beach ridges. see caption for full description. fig 24: photos of holocene beach ridges. see caption for full description. fig 25: geological profile g–h. see caption for full description. fig 26: photo of exposed elevated chalk sheets. see caption for full description. fig 27: historical illustration of møns klint profile. see caption for full description. fig 28: the structure of hvideklint. see caption for full description. fig 29: photos of folding at hvideklint. see caption for full description. fig 30: photo of thrust fault at hvideklint. see caption for full description. fig 31: photo of thrust fault at hvideklint. see caption for full description. fig 32: photo of deformed chalk at hvideklint. see caption for full description. fig 33: photo of folded and thrust-displaced layers at hvideklint. see caption for full description. fig 34: photo of cretaceous chalk at hvideklint. see caption for full description. fig 35: photo of steeply dipping chalk sheets. see caption for full description. fig 38: photo of brecciated clasts of chalk. see caption for full description. fig 39: photo of mud diapirism at hvideklint. see caption for full description. fig 40: conceptual model of glaciodynamic sequence stratigraphy. see caption for full description. fig 41: block diagrams showing the formation of møn’s hilly landscape. see caption for full descript fig 42: structural evolution of møns klint. see caption for full description. fig 43: formation of aborrebjerg and vitmunds nakke. see caption for full description. fig 44: geomorphological elevation model of møn. see caption for full description. fig 45: photo of the boundary between elevated høje møn and the lower, undulating till plain. see ca fig 46: photos of the form and structure of præstebjerg hill. see caption for full description. fig 47: model for the formation of præstebjerg. see caption for full description. fig 48: photo of the 2007 landslide at store taler. see caption for full description. fig 49: photo of fossil hunters following a rockfall. see caption for full description. fig 50: schematic showing erosion and retreat of møns klint. see caption for full description. fig 51: photo of landslide remnants. see caption for full description. fig 52: photo of landslide at abildgaards fald. see caption for full description. table 1: record of large landslides at møns klint between 1801 and 2015. see caption for full descri geological survey of denmark and greenland bulletin 38, 2017, 1-8 1 geological survey of denmark and greenland bulletin 38 • 2017 review of survey activities 2016 edited by adam a. garde and ole bennike geological survey of denmark and greenland danish ministry of energy, utilities and climate 22 geological survey of denmark and greenland bulletin 38 keywords geological survey of denmark and greenland, survey organisations, current research, denmark, greenland. cover photographs from left to right 1. the traditional spear auger remains a principal tool in mapping danish quaternary deposits, but can now be combined with short-range electromagnetic methods for specific purposes such as 3d delineation of contaminated landfill sites. photograph: jakob lautrup 2. elevation map around the stream of hagens møllebæk in north-west jylland with the positions of measuring stations to determine the nitrate runoff from agricultural fields at different times of the year. 3. the monitoring of the surface conditions of the greenland ice sheet by geus is an important tool fo understanding global climate change. photograph: baptiste vandecrux. 4. rescuing primary seismic data from the north sea obtained by private exploration companies to a national electronic database at geus. photograph: peter warna-moors. frontispiece: facing page surface of qagssimiut ice lobe ablation area, southern greenland ice sheet, after a period of extended clear sky conditions that enabled a strong dark ice algal bloom registered in satellite imagery presented in box et al. (2017, this volume). aerial oblique photo near 61°11.006́ n, 46°42.333´w, 820 m elevation. photo: jason e. box, 23 august 2014. chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, department of geoscience, aarhus university; minik rosing, geological museum, university of copenhagen; finn surlyk, department of geosciences and natural resource management, university of copenhagen scientific editors: adam a. garde and ole bennike editorial secretary: jane holst referees (numbers refer to first page of reviewed article): nicolaj k. larsen & lars nielsen, dk (9); anonymous & ole silkjær, dk (13); peter k. engesgaard & bo v. iversen, dk (17); john korstgård & lars nielsen, dk (21); thomas balstrøm & nicolaj k. larsen, dk (25), niels balling & reinhard kirsch, dk (29); anonymous & allan mahler, dk (33); mikael erlström, se & maciej i. kotarba, pl (37); jakob keiding, no & vesa nykänen, fi (41); christopher harrison, ca & martin sønderholm, dk (45); michael k. engkilde & ida l. fabricius, dk (49); xavier fettwis, be & teruo aoki, jp (53); james h. lever, us & christopher j. ries, dk (57); anonymous & jacob c. yde, no (61); asger k. pedersen, dk & daniel sopher, se (65); christian holmegaard, dk & martin l. nayembil, uk (69); jens havskov, no & christopher j. ries, dk (73). illustrations: benny m. schark, allan lindy, stefan sølberg and susanne rømer layout and graphic production: jane holst and carsten e. thuesen printer: rosendahls-schultz grafisk a/s, albertslund, denmark manuscripts received: 12 january – 28 march 2017 final versions approved: 15 march – 19 may 2017 printed: 31 july 2017 issn (print) 1604-8156, isbn (print) 978-87-7871-469-5 issn (online) 1904-4666, isbn (online) 978-87-7871-470-1 citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 38, 76 pp. available from geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk and at www.geus.dk/publications/bull © de nationale geologiske undersøgelser for danmark og grønland (geus), 2017 for the full text of the geus copyright clause, please refer to www.geus.dk/publications/bull 3 44 25 karst sinkhole mapping using gis and digital terrain models p.b. sørensen, h. lykke-andersen, p. gravesen and b. nilsson 29 towards a geothermal exploration well in the gassum formation in copenhagen h. vosgerau, u. gregersen, l. kristensen, s. lindström, a. mathiesen, c.m. nielsen, m. olivarius and l.h. nielsen 33 pre-drilling geothermal assessment of porosity and permeability of the bunter sandstone formation, onshore denmark m.l. hjuler and l. kristensen 37 generation and origin of natural gas in lower palaeozoic shales from southern sweden n.h. schovsbo and a.t. nielsen thailand bangladesh vietnam cambodia greenland kenya ethiopia nigeria cameroon malawi tanzania denmark swedeniceland russia ghana 7 review of survey activities 2016 f.g. christiansen 9 optimising geological mapping of glacial deposits using high-resolution electromagnetic induction data k.e.s. klint, i. møller, p.k. maurya and a.v. christiansen 13 buried valleys in denmark and their impact on the subsurface geological architecture p.b.e. sandersen and f. jørgensen 17 nitrate transport pathways in riparian zones of the hagens møllebæk catchment, northern denmark b. nilsson, a.l. højberg and p. jensen 21 structures and stratigraphy of danian limestone, eastern sjælland, denmark p.r. jakobsen, m.m. rohde and e. sheldon 5 dark grey indicates countries where geus has ongoing or recently completed projects. orange indicates countries with geus projects described in this volume. 41 prospectivity mapping for orogenic gold in southeast greenland b.h. heincke and b.m. stensgaard 45 inversion structures as potential petroleum exploration targets on nuussuaq and northern disko, onshore west greenland e.v. sørensen, j.r. hopper, g.k. pedersen, h. nøhr-hansen, p. guarnieri, a.k. pedersen and f.g. christiansen 49 potential hydrocarbon reservoirs of albian– paleocene age in the nuussuaq basin, west greenland m.l. hjuler, n.h. schovsbo, g.k. pedersen and j.r. hopper 53 greenland, canadian and icelandic land-ice albedo grids (2000–2016) j.e. box, d. van as, k. ste�en and the promice project team 57 new programme for climate monitoring at camp century, greenland w. colgan, s.b. andersen, d. van as, j.e. box and s. gregersen 61 asynchronous ice-sheet development along the central east greenland margin: a glanam project contribution l.f. pérez and t. nielsen 65 �e rescue of seismic �eld data from exploration activities in the danish north sea m.m. hansen and n. rinds 69 an integrated public database for geology, groundwater and drinking water in denmark m. hansen and c.t. �omsen 73 arctic geopolitics and the beginning of earthquake monitoring in denmark and greenland a.l.l. jacobsen thailand bangladesh vietnam cambodia greenland kenya ethiopia nigeria cameroon malawi tanzania denmark swedeniceland russia ghana 66 7 review of survey activities 2016 flemming g. christiansen deputy director 2016 was a year of transition in the international geological society. like previous years, it was still a period with lots of reorganisation and cut-backs in the energy and mineral industry but also with obvious signs of an even more complex business pattern where well-organised geological data and easy access to geological knowledge will be strongly needed. compared to the industry, and also to some of the authorities that are regulating the use of natural resources, geus has a high degree of continuity of both activities and personnel; in these years with strong focus on key strategic goals. this creates and develops a corporate memory that can be easily applied even with sudden shifts of a political or business agenda. this also establishes geus as the starting point for getting access to geological information from denmark and greenland, and makes geus a natural collaboration partner nationally as well as internationally. publication of results, both internationally and in our own series is a key factor for keeping a high scientific standard and for documentation of this towards society, authorities and industry. this issue of geus’ review of survey activities shows like in previous years a broad spectrum with a total of 17 papers covering many different activities. the papers give a good overview in their own right but with systematic references to background reports and other papers, they also provide an easy way to dig deep into the important work that geus is carrying out. eight papers are on denmark, six on greenland and three on other themes. activities in denmark geus works on many different – and often closely interrelated – topics in denmark such as the use of water, energy, and mineral resources, protection of nature under significant climatic changes, and by making up-to-date geological and geophysical data and information easily available for all sorts of purposes. the use of groundwater is very important for denmark, and geus carries out many studies on water resources and their protection in connection with climate changes, environmental impact and domestic use. systematic geological and geophysical studies provide not only specific water data, but also give a much better understanding of the geological models that can be used to predict resources and regulate their use. one paper with a case study from samsø describes how traditional geological mapping can be optimised by using high-resolution electromagnetic induction data. another paper gives an overview of buried valleys in denmark, their geological architecture and very significant length, which has great implications for the groundwater resources in many parts of denmark. a third paper gives a detailed description of nitrate transport pathways in one catchment area in northern jylland where results from year-around monitoring can be used for planning and regulating to meet the demands of the eu water framework directive. in many areas in denmark, chalk and limestone are directly exposed or found near terrain surface with only a thin cover of quaternary deposits. such deposits are important as a mineral resource for many different purposes and due to their groundwater resources and they may also control structures in overlying sediments. one paper documents the structures and stratigraphy of danian limestone on eastern sjælland, whereas another study uses gis and digital terrain models to map karst sinkholes in areas of jylland and to interpret the geological and climatic conditions and anthropogenic activities that influence their development. denmark has a large potential for subsurface geothermal energy and for heat and energy storage. following the recent launching of the geothermal webgis portal with all geological and geophysical data relevant for geothermal exploration, geus has continued with more detailed studies of specific areas and on specific parameters. one paper gives a summary of a eudp-supported project on the planning of exploration in urban areas of copenhagen by targeting on shallower reservoirs in the gassum formation and by using smaller drilling rigs. another paper has focus on the porosity and permeability variation of the bunter sandstone formation that is a potential geothermal reservoir in large onshore areas of southern denmark. shale-gas production has been a major game changer in the energy sector for quite some years, and we have seen © 2017 geus. geological survey of denmark and greenland bulletin 38, 7–8. open access: www.geus.dk/publications/bull 88 many studies of the resource potential and some exploration in europe, including denmark and southern sweden. one paper gives a detailed description of the gas composition and discuss biogenic versus thermogenic origin of gas in lower palaeozoic shales from scientific core holes in southern sweden. activities in greenland once again there was a high level of geological and glaciological activities in greenland in 2016, both traditional studies with focus on the mineral and petroleum potential and exploration studies as well as monitoring and research related to climate changes and their effect. geus has been active with regional mapping and research in south-east greenland for quite some years, and some of the geo-datasets are important for identifying exploration targets. one paper argues for the possibility of orogenic gold deposits by using aeromagnetic and stream sediment geochemistry data in a minerals systems model. a large number of data and proxies maps have been applied to construct a final prospectivity map pointing towards the most interesting possibilities in the tasiilaq area. the nuussuaq basin in west greenland has served as an analogue geological model for offshore petroleum exploration for decades but there may also be an onshore potential. as preparation for both offand onshore licensing rounds, geus has compiled key data in a gis model and re-studied especially the large-scale structures in more detail using photogrammetric mapping. one paper describes a very large inversion anticline on central nuussuaq that could be an interesting drilling target. a revised migration pathway model is proposed; this could explain distribution of both oil and gas seeps outside the traditional area with common seepage. another paper describes the reservoir properties of quite a number of sedimentary and volcanic successions of aptian to paleocene age in the nuussuaq basin. monitoring programmes of the greenland ice sheet and research based on local ground truth data from stations on the ice – and from fjords and nearby offshore areas – are very important contributions from geus to global climate models. one paper gives an example of albedo grids measured from satellite and checked with local data to give a much better de-noising and bias correlation. another paper outlines a new climate monitoring and radar mapping programme in the camp century area in northern greenland that will be managed by geus; some historical information and background models are also presented. a third paper describes the ice-sheet development along the central east greenland margin in late miocene to recent time where thick and different types of glacigenic deposits are interpreted using seismic data. other themes as the national geological data center geus has a strong obligation to make all data available to authorities, educational and research institutes and to private enterprises. the ever accelerating changes in technology make this a great challenge in terms of competence, capacity and economic resources, both for collection and quality control of the actual data and for the database and distribution systems. the value for society is, however, very high and there is a strong focus on digitalisation strategy in denmark these years. one paper describes the tremendous effort that was made to rescue most of the seismic field data from exploration activities in the 1980s and 1990s in the danish north sea and how successfully this worked in collaboration with the operators. another paper introduces jupiter, the integrated public database for geology, groundwater and drinking water in denmark, and describes the data systems, agreements and management behind it. finally, the last paper presents the early history of earthquake monitoring in denmark and greenland, and discusses how this was as geopolitically important more than a century ago as it is today. data article | short tegner et al. 2023: geus bulletin 53. 8316. https://doi.org/10.34194/geusb.v53.8316 1 of 8 a whole-rock data set for the skaergaard intrusion, east greenland christian tegner*1 ,lars peter salmonsen2 , marian b. holness3 , charles e. lesher1 , madeleine c. s. humphreys3,4 , peter thy5 , troels f. d. nielsen6 1department of geoscience, aarhus university, aarhus, denmark; 2rambøll, aarhus, denmark; 3department of earth sciences, university of cambridge, cambridge, uk; 4department of earth sciences, durham university, durham, uk; 5department of earth and planetary sciences, university of california, davis, usa; 6department of mapping and mineral resources, geological survey of denmark and greenland, copenhagen, denmark abstract we report a compilation of new and published whole-rock major and trace element analyses for 646 samples of the skaergaard intrusion, east greenland. the samples were collected in 14 stratigraphic profiles either from accessible and well-exposed surface areas or from drill core, and they cover most regions of the intrusion. this includes the layered series, the upper border series, the marginal border series and the sandwich horizon. the geochemical data were obtained by a combination of x-ray fluorescence and inductively coupled plasma mass spectrometry. this data set can, for example, be used to constrain processes of igneous differentiation and ore formation. *correspondence: christian.tegner@geo.au.dk received: 05 apr 2022 revised: 01 mar 2023 accepted: 13 apr 2023 published: 15 june 2023 keywords: skaergaard intrusion, layered mafic intrusion, bulk-rock geochemical data, x-ray fluorescence (xrf), inductively coupled plasma mass spectrometry (icp-ms) abbreviations a.s.l.: above sea level f: mass fraction of melt remaining in the chamber geus: geological survey of denmark and greenland. hz: hidden zone icp-ms: inductively coupled plasma mass spectrometry loi: mass lost on ignition ls: layered series lz: lower zone mz: middle zone s.d.: standard deviation sh: sandwich horizon uz: upper zone xrf: x-ray fluorescence geus bulletin (eissn: 2597–2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: kerstin saalmann (geological survey of norway) reviewed by: rais latypov (university of witwatersrand, south africa), howard naslund (binghamton university, usa) funding: see page 7 competing interests: see page 7 additional files: see page 7 tabular abstract geographical coverage the skaergaard intrusion occupies a c. 11 × 8 km outcrop area of layered gabbroic rocks at uttentals sund, kangerlussuaq area, east greenland. located at c. 68°9 ′n and 31°41′ w. temporal coverage palaeogene (c. 56.0 ma) subject(s) cosmochemistry and geochronology, economic geology, geochemistry, igneous rocks and processes data format(s) major and trace element compositions reported in an excel spreadsheet. sample collection & analysis samples (n = 646) taken from surface outcrops and drill cores were analysed by x-ray fluorescence and inductively coupled plasma mass spectrometry (icp-ms). the samples are stored and curated at: aarhus university (surface samples from the layered series and upper border series); geological survey of denmark and greenland (surface samples from the layered series); natural history museum of denmark, university of copenhagen (drill core samples) and the harker collection of the sedgwick museum, university of cambridge (cambridge 1966 drill core and surface samples of the marginal border series). parameters major and trace element whole-rock compositions. related publications: tegner 1997; tegner et al. 2009; salmonsen & tegner 2013; holness et al. 2015, 2017, 2022; thy et al. in press. potential application(s) for these data this data set can, for example, be used to constrain processes of igneous differentiation and ore formation. data collection to examine the petrology and ore bodies of the skaergaard intrusion, east greenland, we have collected hundreds of samples during six field expeditions between 1993 and 2017. in addition, we have collected samples from drill core material housed at the natural history museum of denmark https://doi.org/10.34194/geusb.v53.8316 https://orcid.org/0000-0003-1407-7298 https://orcid.org/0009-0008-8005-0150 https://orcid.org/0000-0001-9911-8292 https://orcid.org/0000-0003-4033-4809 https://orcid.org/0000-0001-9343-1237 https://orcid.org/0000-0002-9267-5798 https://orcid.org/0000-0002-4932-3869 mailto:christian.tegner@geo.au.dk tegner et al. 2023: geus bulletin 53. 8316. https://doi.org/10.34194/geusb.v53.8316 2 of 8 www.geusbul let in.org (university of copenhagen) and the harker collection of the sedgwick museum (university of cambridge). here, we report a compilation of 646 whole-rock analyses for these samples attached as one excel spreadsheet (supplementary data file 1). all samples were analysed by x-ray fluorescence (xrf); most of these data (n = 409) were published in tegner (1997), tegner et al. (2009), salmonsen & tegner (2013), holness et al. (2015, 2022) and thy et al. (in press). the remaining analyses (n = 237) are reported here for the first time apart from p2o5 data (n = 167), which were reported in holness et al. (2017). a subset of 271 samples were analysed by inductively coupled plasma mass spectrometry (icp-ms). about half of these (n = 130) were published in tegner et al. (2009) and thy et al. (in press); the remaining analyses (n = 141) are reported here for the first time. the analysed samples mainly represent mafic cumulate rocks (n = 623) but also include gabbropegmatite and granophyric pods and layers (n = 23). details of subsets of the present bulkrock data set have been described and discussed in a number of publications (e.g. tegner 1997; tegner et al. 2009; thy et al. 2009, in press; tegner & cawthorn 2010; mckenzie 2011; salmonsen & tegner 2013; namur et al. 2013, 2014; holness et al. 2015, 2017, 2022; nielsen et al. 2015; keays & tegner 2016; pedersen et al. 2021). the samples were collected in 14 stratigraphic profiles as shown on the geological map (fig. 1), in a schematic cross-section (fig. 2), listed in table 1 and described in detail in supplementary data file 2. the sample sections thus cover the known stratigraphy and rock units reported in the layered series (ls: including hidden zone, hz; lower zone a, lza; lower zone b, lzb; lower zone c, lzc; middle zone, mz; upper zone a; uza; upper zone b, uzb; upper zone c, uzc), the marginal border series (mbs: including lower zone a*, lza*; lower zone b*, lzb*; lower zone c*, lzc*; middle zone*, mz*; upper zone a*; uza*; upper zone b*, uzb*; upper zone c*, uzc*), the upper border series (ubs: including lower zone a´, lza´; lower zone b´, lzb´; lower zone c´, lzc´; middle zone´, mz´; upper zone a´; uza´; upper zone b´, uzb´; upper zone c´, uzc´), and the sandwich horizon (sh). this zonal subdivision is outlined in figs 1 and 2. a field photograph (fig. 3) shows an example of layered rocks (layered series). further selected outcrops are shown in fig. s1 (supplementary data file 2). previous work has shown that the intrusion represents the result of prolonged, uninterrupted differentiation of a tholeiitic magma (wager & deer 1939; wager & brown 1968; naslund 1984; hoover 1989; mcbirney 1996; irvine et al. 1998). palladium and gold mineralisations have been identified in the ls (andersen et al. 1998; nielsen et al. 2015). a key point is that the ls, mbs and ubs appear to represent continuous and synchronous crystallisation on the floor, margins and roof and sh the products of the most evolved, last drops of magma in the interior of the intrusion. these rocks, therefore, allow evaluation of the processes resulting in igneous differentiation and ore formation in opposite positions relative to gravity (e.g. mcbirney 1995). sample profiles the stratigraphic profiles (n = 14) are summarised in table 1 and illustrated in figs 1 and 2. the profiles cover most regions of the intrusion and were collected either from accessible and well exposed surface areas or from drill core. the details of the sample profiles are described in supplementary data file 2. sampling strategy the sampling was directed to obtain mainly average rock compositions in systematic stratigraphic sections. additional samples of outcrop features such as gabbropegmatites, subzone boundaries and layered structures were also included, for example, the ‘wavy pyroxene rocks’ and colloform banding of the mbs (humphreys & holness 2010; namur et al. 2013). the sample positions were recorded by gps and altimeter readings (supplementary data file 1). for ls and ubs, the stratigraphic thicknesses were calculated relative to the local strike and dip of layering as described previously (tegner et al. 2009; salmonsen & tegner 2013) and are listed in supplementary data file 1. within the limitations of outcrops, we aimed to sample at regular stratigraphic intervals. for ls and ubs, the average stratigraphic interval was 12 ± 13 m (1 s.d.). for mbs, the lateral distance from the contact is recorded and the average spacing between samples was 18 ± 6 m (1 s.d.; holness et al. 2022). calculation of fraction of melt remaining (f) the box-like appearance of the intrusion with onionring distribution of zones and subzones (fig. 2) implies that stratigraphic thickness, and mass proportions are not proportional (nielsen 2004). based on mass proportions estimated for each subzone in the floor, wall and roof series (nielsen 2004), the mass fraction of melt remaining in the chamber, f, can be estimated for subzone boundaries. for the ‘reference profile’ of ls, a second-order polynomial was fitted to subzone boundaries to relate f to stratigraphic height (h) and given below as equation 1 (tegner et al. 2009): f = 1.091 × 10-7h2 – 5.9064 × 10-4h + 0.7678 (1) the f values for the 90-10 and 90-23 drill core samples were also estimated using equation 1 and tied to the ‘reference profile’ at the h of the uza/b boundary https://doi.org/10.34194/geusb.v53.8316 http://www.geusbulletin.org tegner et al. 2023: geus bulletin 53. 8316. https://doi.org/10.34194/geusb.v53.8316 3 of 8 www.geusbul let in.org fig. 1 geological map of the skaergaard intrusion and adjacent host rocks. the rocks that solidified at the floor (layered series composed of lower zone, lz, middle zone, mz, and upper zone, uz), walls (marginal border series, mbs) and roof (upper border series, ubs) are shown. also shown are the approximate locations of 14 sample profiles (surface samples and drill cores) that are studied here. the sample profiles are numbered 1 to 14 and listed in table 1. further subzone abbreviations are in the text. modified from mcbirney (1989). uttental sund forbindelsesgletscher lzb lzc lzc lza mbs uza 2 km uza uzc uzb uzbmz mz uttental plateau kraemer ø figure 1 dobbelt gletscher skaergaard bay hammer gletscher basistoppen lz mz uz mbs ubs basalt contact of the skaergaard intrusion la ye re d se rie s gneiss sea ice later mafic intrusion drill core sample profile [1] [1] [2] [1] [1] [4] [1] [3] [5][6] [7] [8] [9] [14] [13] [12] [11] [10] 68 °0 8ʹ n 68 °1 3ʹ n 31°45ʹw 31°35ʹw https://doi.org/10.34194/geusb.v53.8316 http://www.geusbulletin.org tegner et al. 2023: geus bulletin 53. 8316. https://doi.org/10.34194/geusb.v53.8316 4 of 8 www.geusbul let in.org (1615 m). similarly, the f values for the samples of the cambridge core were estimated with equation 1, setting h to zero at the lza/hz boundary (holness et al. 2015). for ubs, f values of subzone boundaries were fixed to the same values as for ls, and f values of samples were related to stratigraphic height by linear interpolation (salmonsen & tegner 2013). similarly for mbs, the f values at subzone boundaries were assumed equal to those of ls, and the f values of the samples were related to the distance between subzone boundaries by linear interpolation. finally, in the sections crossing the sh, we assigned an f value of zero to the sample with the lowest mgo content (≤0.02 wt%). the estimated f values are reported in supplementary data file 1. analytical methods all samples were collected, prepared and analysed in the same way. from surface outcrops, we collected samples weighing 1–4 kg and avoiding alteration veins. the samples were trimmed for surface weathering by sawing, and an aliquot (100–400 g) was crushed to small aggregates (<2 cm) in a hydraulic steel press. this was followed by splitting, pre-contamination of a corundum shatterbox, cleaning and finally powdering of c. 30 g. the drill core material was prepared in the same way with one exception. the powders of drill core 90-22 (n = 51) were prepared using a steel jaw crusher and a tungsten-carbide shatterbox. all samples were prepared at aarhus university as described in tegner et al. (2009). fig. 2 schematic cross-section of the skaergaard intrusion showing the distribution of rocks that solidified at the floor (layered series), the walls (marginal border series, mbs) and the roof (upper border series, ubs). also shown are the approximate locations of 14 sample profiles (surface samples: blue lines, drill cores: orange lines) that are studied here. the sample profiles are also shown on the map in fig. 1 and numbered 1 to 14 as listed in table 1. the floor, wall and roof sequences are divided into subzones (hz–uzc) depending on the appearance and disappearance of primary (cumulus) crystal phases as marked on the subzone boundaries. abbreviations: hz: hidden zone. lz: lower zone. mz: middle zone. uz: upper zone. sh: sandwich horizon. further subzone abbreviations and nomenclature are in the text. modified from mcbirney (2002) and nielsen (2004). 2.5 2.0 1.5 1.0 0.5 0 lza lzb lzc mz uza uzb uzc lza’ lzb’ lzc’ mz’ uza’ uzb’ uzc’ hz lza* lzb* lzc* m z* u za* u zb* st ra tig ra ph ic th ic kn es s (k m ) sandwich horizon platinova reefs 3.0 east west c. 8 km aug+ mt+ ap+ olol+ granophyre uzb’+c’ [1] [6] [1] [1] [1] [1] [11] [10] [8] [7] [9] [5] [4] [3] [2] [14] [13] [12] mbs mbs ubs layered series basistoppen sill melanogranophyre https://doi.org/10.34194/geusb.v53.8316 http://www.geusbulletin.org tegner et al. 2023: geus bulletin 53. 8316. https://doi.org/10.34194/geusb.v53.8316 5 of 8 www.geusbul let in.org ta bl e 1  o ve rv ie w o f s am pl es a nd s am pl e pr ofi le s sa m pl e pr ofi le ro ck s er ie s su bz on es n o. o f s am pl es n o. o f s am pl es n o. o f s am pl es n o. o f s am pl es re fe re nc e n am e n o. b     to ta l cu m ul at e m el an og ra no ph yr e g ab br op eg m at ite re fe re nc e pr ofi le a [1 ] la ye re d se ri es lz a, l zb , l zc , m z, u za , u zb , u zc 13 8 13 5 2 1 te gn er (1 99 7) ; t eg ne r et a l. (2 00 9) ; t hy e t a l. (in p re ss ) ca m br id ge d ri ll co re [2 ] la ye re d se ri es h z, l za 12 1 12 1 h ol ne ss e t a l. (2 01 5) d ob be lt g le ts ch er [3 ] la ye re d se ri es h z, l za 8 8 th is s tu dy 90 -1 0 dr ill c or e [4 ] la ye re d se ri es m z, u za , u zb 81 81 th is s tu dy 90 -2 3 dr ill c or e [5 ] la ye re d se ri es u za , u zb , u zc 87 87 th is s tu dy n w b as is to pp en [6 ] la ye re d se ri es /u pp er bo rd er s er ie s u zc , s h , u zc ´ 21 12 9 th is s tu dy ki le n [7 ] u pp er b or de r se ri es / la ye re d se ri es h z´ , l za ´, l zb ´, l zc ´, m z´ , u za ´, u zb ´, u zc ´, s h ; u zc 33 26 6 1 sa lm on se n & t eg ne r (2 01 3) h am m er p as s [8 ] u pp er b or de r se ri es h z´ , l za ´, l zb ´, l zc ´, m z´ , u za ´ 25 25 sa lm on se n & t eg ne r (2 01 3) br ød re to pp en [9 ] u pp er b or de r se ri es h z´ , l za ´, l zb ´, l zc ´, m z´ , u za ´, u zb ´, u zc ´, s h 37 34 3 sa lm on se n & t eg ne r (2 01 3) sy dt op pe n [1 0] u pp er b or de r se ri es h z´ 11 11 sa lm on se n & t eg ne r (2 01 3) sk ae rg aa rd b ay [1 1] u pp er b or de r se ri es h z´ , l za ´, l zb ´ 15 15 th is s tu dy sk ae rg aa rd p en in su la [1 2] m ar gi na l b or de r se ri es lz a* , l zb *, l zc *, m z* , u za *, u zb * 33 33 h ol ne ss e t a l. (2 02 2) iv na rm iu t i sl an d [1 3] m ar gi na l b or de r se ri es lz a* , l zb *, m z* 17 17 h ol ne ss e t a l. (2 02 2) kr ae m er ø [1 4] m ar gi na l b or de r se ri es lz a* , l zb * 19 19 h ol ne ss e t a l. (2 02 2) to ta l n o. o f s am pl es 64 6 62 4 20 2   a r ef er en ce p ro fil e co ns is ts o f s ur fa ce s am pl es fr om u tt en ta l p la te au , k ra em er ø , p uk ug ag ry gg en /f or bi nd el se sg le ts ch er , w es t b as is to pp en , a nd d ri ll co re 9 022 . b n um be rs la be lle d in f ig . 1 . lz : l ow er z on e. m z: m id dl e zo ne . u z: u pp er z on e. h z: h id de n zo ne . s h : s an dw ic h h or iz on . https://doi.org/10.34194/geusb.v53.8316 http://www.geusbulletin.org tegner et al. 2023: geus bulletin 53. 8316. https://doi.org/10.34194/geusb.v53.8316 6 of 8 www.geusbul let in.org the major and trace element data were obtained by a combination of xrf at aarhus university and icp-ms at the university of california, davis and acmelabs as described in tegner et al. (2009), thy et al. (in press) and tegner et al. (2019), respectively. concentrations of feo were determined by titration with potassium dichromate. the mass lost on ignition (loi) was determined by heating the powder in air in a muffle furnace at 950°c for 3 h. the values obtained for certified reference materials, bhvo-1 and bir-1, are reported in supplementary data file 3. xrf analyses of bhvo-1 and bir-1 (n = 53–63) demonstrate that the relative variation of repeated analyses is less than 4.5% (1 s.d./average value) for most major element oxides. however, in bir-1, the relative variation is higher (14%) for k2o, which has a relatively low concentration (0.027 wt%; jochum et al. 2016). for the trace elements measured by xrf, the relative variation of the transition metals (v, cr, ni, cu, zn) and sr are within 5%. in bhvo-1, the relative variation is moderate for rb, y, zr, nb and ba (4–12%) and higher for ce (20%). in bir-1, which is depleted in these elements relative to bhvo-1 (jochum et al. 2016), the relative variation of repeat analyses is somewhat higher for rb, y, zr and nb and pb (7–28%) and much higher for ba and ce. the accuracy or relative deviation from the preferred values for bhvo-1 is within 11% for all oxides and trace elements. similar values were obtained for the accuracy of bir-1, except for rb, nb and ce, which have sub-ppm preferred values. in conclusion, the xrf data can generally be viewed as accurate down to a few ppm. repeat icp-ms analyses of standards at acmelab (n = 14) and university of california (n = 6) deviate less than 7 and 18%, respectively, from the preferred values for all trace elements reported in this study (supplementary data file 3). data description and main features the bulk compositions of cumulate rocks, such as those reported here, represent a mix of accumulated liquidus crystals (cumulus) and interstitial material (intercumulus) derived from crystallisation of interstitial melt (wager et al. 1960; irvine 1982). the present data set thus tracks changes in the compositions fig. 3 field photo showing the northwards view from kraemer ø. metre-scale modal layering occurs in the middle zone (mz) (foreground, lower two-thirds) and in the background where the triple group can be seen on wager peak (c. 1200 m a.s.l.). wager peak triple group senw mz mz fig. 4 example compositional data available in the data set. a: whole-rock feototal versus sio2 (wt%) for the skaergaard intrusion. one outlier at 17 wt% sio2 and 50 wt% feo is not shown. b, c: stratigraphic variations in whole-rock feototal and rb contents. data from supplementary data file 1. 0 5 10 15 20 25 30 35 40 20 30 40 50 60 70 80 fe o (w t% ) sio2 (wt%) hz+lz mzls ubs mbs uz lz* mz* uz* hz´+lz´ mz´ uz´ gabbropegmatite melanogranophyre 0 5 10 15 20 25 30 35 40 fe o (w t% ) 10 20 30 40 50 0.00.20.40.60.81.0 r b (p pm ) fraction of magma remaining (f) 60 hz lza lzb lzc mz uza uzb uzc (b) (a) (c) https://doi.org/10.34194/geusb.v53.8316 http://www.geusbulletin.org tegner et al. 2023: geus bulletin 53. 8316. https://doi.org/10.34194/geusb.v53.8316 7 of 8 www.geusbul let in.org and proportions of minerals and trapped melt during solidification of the skaergaard magma chamber. importantly, the bulk-rock compositions do not directly represent liquid compositions. the data set can therefore be used to evaluate igneous processes during crystallisation and ore formation. figure 4a, for example, shows that bulk-rock feototal and sio2 vary considerably and display a negative correlation. these two oxides also show systematic variations between zones (hz, lz, mz, uz). the compositions generally overlap between ls, mbs and ubs rocks although the most feo-rich rocks occur in ls. in the uz equivalents, the ubs rocks are enriched in sio2 relative to ls and mbs rocks. not surprisingly, the highest sio2 and the lowest feo values are seen for granophyres sensu lato. figure 4 also shows two examples of stratigraphic variations plotted against the calculated fraction of melt remaining (f). in fig. 4b, feo generally increases from lz to uz equivalents and displays a marked increase across the lzb/lzc boundary, reflecting accumulation of magnetite and ilmenite. in the lower and middle part of the stratigraphy (hz–uza), the feo contents are comparable in the floor, wall and roof rocks. however, in uzb’ and uzc’ of the roof (ubs), feo is markedly lower compared to ls and mbs rocks. figure 4c shows the stratigraphic trends of the incompatible element rb. in the lower and middle parts (hz–uza), the trends are relatively flat and display comparable values in rocks from ls and mbs, while higher values are found in ubs rocks. closer to sh (uzb and uzc equivalents), rb increases exponentially in the cumulate rocks and shows the highest values in the melanogranophyres. the compiled whole-rock data set can, for example, be used to constrain processes of igneous differentiation and ore formation. acknowledgements we are grateful to the danish lithosphere centre for supporting field work in 2000. platina resources were accommodating during the 2008 and 2011 field seasons. the geological survey of denmark and greenland (geus) helped with field logistics in 2017. platinova resources ltd. are thanked for access to drill core material. we thank c. kent brooks for inspiring this project. we are indebted to jakob k. keiding for help with field work, sample preparation and discussion. we also enjoyed assistance and company in the field from jens c.ø. andersen, olivier namur, anja k.m. fonseca and joel a. simpson. sidsel grundvig, ingrid aaes and jette villesen, aarhus university, are thanked for help with sample preparation and x-ray fluorescence analyses. we thank two reviewers, rais latypov and howard naslund, as well as kerstin saalmann for careful editorial handling. additional information funding statement this work was supported by funding from danish national science research council (ct, tfdn), the danish national research foundation (tfdn, cel, ct), the carlsberg foundation (ct), aarhus university (ct, lps), the uk natural environment research council (mbh, mcsh), the us national science foundation under grant number nsf-ear-0208075 (cel), the uk royal society international joint project (mbh, ct). author contributions ct: conceptualisation, data curation, funding acquisition, investigation, methodology, supervision, visualisation, writing – original draft. lps: investigation, methodology, writing – review and editing. mbh: conceptualisation, data curation, funding acquisition, investigation, methodology, writing – review and editing. cel: investigation, conceptualisation, funding acquisition, methodology, writing – review and editing. mcsh: investigation, methodology, writing – review and editing. pt: investigation, methodology, writing – review and editing. tfdn: investigation, conceptualisation, funding acquisition, methodology, writing – review and editing. competing interests the authors declare no competing interests. additional 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https://doi.org/10.1093/petrology/egp020 https://doi.org/10.1093/petrology/egp020 https://doi.org/10.1007/s00410-008-0361-6 https://doi.org/10.34196.v53.8327 https://doi.org/10.34196.v53.8327 https://doi.org/10.1093/petrology/1.1.73 https://doi.org/10.1093/petrology/1.1.73 geological survey of denmark and greenland bulletin 17, 2009, 37-40 the danish landscape is characterised by low relief and consequently the risk of soil erosion is low compared to many central and southern european countries with more variable terrain (european environment agency 2000; van der knijff et al. 2000). however, even in countries with less intensive erosion, water-induced soil erosion is recognised as an increasingly important environmental issue due to its role in the transport of nutrients, pesticides and other contaminants to rivers, lakes and coastal waters (e.g. stone 2000). there are relatively few monitoring programmes of sediment erosion in denmark and those that exist typically only cover the last few decades (veihe et al. 2003). therefore, it is difficult to identify long-term trends and baseline conditions. however, the geological record provides insight into patterns and rates of soil erosion through time. lake and fjord sediment archives are especially useful because they can provide continuous and undisturbed sediment successions that can be ana lysed at high temporal resolution. these sediment records can be examined using a variety of approaches that include sedimentological, geochemical and biological analyses and thus generate important insight into baseline states, trajectories and responses to forcing mechanisms over long timescales, which would otherwise be difficult to obtain. the records are typically limited to yielding insight only into the average rate of erosion and they often provide little information about the spatial extent or distribution of erosional processes within a catchment area. in this paper we present an example of a long-term erosional record from gudme sø, fyn (fig. 1). numerous archae ological excavations have been undertaken near gudme sø, and the geological survey of denmark and greenland has carried out analyses of sediments from the lake to provide a continuous picture of landscape changes. sediment accumulation rates of minerogenic matter and pollen analysis are here used to ex plore the erosional response to changes in land-use. study site and methods gudme sø is located on south-east fyn c. 5 km from the sea in a gently undulating terrain with surface deposits consisting of clay till (fig. 1). the lake has a surface area of 9 ha, a mean water depth of 0.5 m, and a catchment area of 48 ha. it has no major natural inlets or outlets. an 11.8 m long sediment core of slightly humic gyttja with low calcareous content and varying silt and clay contents was retrieved from the lake. age determination was provided by radiocarbon dating, and the chronology at the top of the core was further refined by linear interpolation, using the first occurrence of spheroidal carbonaceous fly-ash particles and the date of the sediment surface (i.e. the year of coring) as bracketing ages (fig. 2). sediment accumulation rates of minerogenic matter (sar-min in mg/cm2 per year) were calculated and used as prox ies of soil erosion from the catchment area to the lake (e.g. mackereth 1966). the changes in sar-min values were compared to pollen-inferred changes in vegetation and landuse from the same sediment core. the increases and decreases in non-arboreal pollen (nap) were used as a record of changes in the proportion of open land. it is assumed that the variation in the sum of taxa of cultivated plants and plantago lanceolata (ribwort plantain) reflects changes in the extent of arable land, and pastures or meadows, respectively (behre 1981; gaillard 2007). the record of colonies of the green algae pediastrum sp. is used as proxy evidence for in-lake productivity (bradshaw et al. 2005). soil erosion and land-use change during the last six millennia recorded in lake sediments of gudme sø, fyn, denmark peter rasmussen and jesper olsen © geus, 2009. geological survey of denmark and greenland bulletin 17, 37–40. available at: www.geus.dk/publications/bull 37 fig. 1. aerial photograph of gudme sø seen from the south-east. the lake size is c. 490 × 250 m. today the lake catchment area (48 ha) consists of 50% tilled, 25.8% natural and 24.2% built-up area. photograph courtesy of the royal danish air force. rosa_2008:rosa-2008 01/07/09 15:48 side 37 results and discussion chronology the 14c dates from the gudme sø sediment core are based on bulk sediment samples and may overestimate their actual ages due to the hard-water effect (björck & wohlfarth 2001). however, the caco3 content of the sediments is very low and the ages of key-pollen, stratigraphical levels at gudme sø are in agreement with other independent ages from corresponding levels in danish bogs where the dated peat is unaffected by hard water (table 1). hence, we consider the gudme sø age–depth model to be reasonably accurate and reliable within c. 100 years. in the following, all ages are given in calendar years bc/ad. incipient landscape disturbance in fig. 3 the history of the last c. 6000 years of soil erosion to the gudme sø basin is compared with contemporary changes in vegetation and land-use. coincident increases in total nap and taxa of cultivated plants and in sar-min suggest a causal link between agricultural activities and soil loss. the low sar-min values in the late mesolithic (c. 4200– 3900 bc) suggest a low level of soil disturbance, which is in good agreement with the pollen data. very low nap values (c. 2%) indicate that the landscape was dominated by closedcanopy forests with only a sparse field layer vegetation (iver sen 1973). during the mesolithic, anthropogenic impact on the vegetation appears to have been negligible, and landscape disturbance was driven by natural agencies, implying that the sediment flux to the lake at this time represents a ‘natural’ baseline state. in denmark the cultural shift from the mesolithic (hunting-fishing-gathering) to the neolithic (farming) way of living took place c. 3900 bc. approximately 200 years later (c. 3650–3500 bc) a marked increase in both sar-min and nap values at gudme sø indicate an erosional event clearly triggered by deforestation that was associated with incipient agricultural activities. at that time the first pollen types unambiguously indicative of crop cultivation and pastoral farming occurred (cereals and plantago lanceolata, respectively). interestingly, the data indicate that catchment disturbance and agricultural activities precede accelerated erosion rates by c. 150–200 years, thus showing a delay or threshold effect in the system’s response. between c. 3350 and 3200 bc soil erosion rates increased again, although this time not correlated with pollen evidence of land clearance or intensified agriculture. low soil erosion rates between c. 3200 and 2650 bc and low nap values indicate catchment stability. at about 2650 bc a renewed and short-lived (c. 2650–2550 bc) deforestation was accompanied by intensified agricultural activities reflected in high percentages of nap, plantago lanceolata and cultivated taxa. despite the apparent intensity of this brief land clearance episode, the erosional response was moderate. this might be 38 20001000010002000300040005000 age (calendar years bc/ad) d ep th b el ow s ed im en t su rf ac e (c m ) k-5839 k-6022 k-5840 k-6023 k-5771 k-6024 k-5770 k-6025 k-5841 k-6026 k-5842 k-5769 k-5768 200 300 400 500 600 700 800 900 1000 1100 1200 100 table 1. 14c dates of pollen-analytical levels in gudme sø and three danish raised bogs gudme fuglsø abkær holmegaard sø mose mose mose data partly from odgaard (1994). 1st centaurea cyanus (cornflower) ad 1320 ad 1360 ad 1240 ad 1350 1st secale cereale (rye) 170 bc 80 bc 80 bc 80 bc fagus sylvatica (beech) 3–5% 960 bc 890 bc 240 bc 920 bc fagus sylvatica (beech) > 0.5–1% 1460 bc 1320 bc 1320 bc 1400 bc 1st plantago lanceolata (ribwort plantain) 3650 bc 3630 bc 3000 bc 3800 bc elm decline 3990 bc 3830 bc 3870 bc 3500 bc fig. 2. age–depth model for the gudme sø sediment succession based on 13 calibrated 14c dates, the date of the first occurrence of spheroidal carbonaceous fly-ash particles (blue dot at ad 1900; odgaard 1993) and the age of the sediment surface (red dot at ad 1989). the k-numbers are the laboratory numbers for the dated samples. the small diagrams show probability distributions of the calibrated ages. the age–depth model was established using the programme bpeat (blaauw & christen 2005). rosa_2008:rosa-2008 01/07/09 15:48 side 38 due to the type of land-use at a time when subsistence economy most likely was predominantly based on pastoral farming. in contrast to arable farming, such land-use implies the maintenance of permanently grass-covered pasture that would be less prone to erosion than cultivated fields (veihe et al. 2003). after 2550 bc, low soil erosion rates indicate a long period of relative stability in the catchment area, which lasted until the middle of the bronze age at c. 1000 bc. landscape transformation for c. 150 years, between 1000 and 850 bc, the landscape around gudme sø changed drastically and became far more open than previously as suggested by the increase in nap. extensive forest clearance and expansion of areas used for livestock grazing and crop cultivation led to comprehensive landscape disturbance and a marked increase in erosion rates to the lake. the contemporary increase in pediastrum sp. indicates a larger nutrient loss from the catchment area to the lake due to soil in-wash, which enhanced the in-lake productivity. after c. 200 years, the sediment influx to the lake gradually declined and stabilised, despite a continued high an thro pogenic impact on the landscape as indicated by the continued high nap values. in the iron age around 100 bc, the erosion rates in creased slightly and remained almost constant until c. ad 400. thereafter, there was a short-lived disturbance period around ad 400–500 with intensified arable farming (in creased abundances of cultivated taxa), accelerated soil erosion, and higher lake productivity (raised percentages of pediastrum sp.). high percentages of cannabis-type pollen in this period (not shown) indicate that the lake was used for retting hemp, a process in which stems are soaked in water to free their bast fibres from the surrounding tissue in order to produce fibres for example for cloth and rope making. hemp-retting in the lake most likely contributed to the increased input of minerogenic material, as plants for retting may have been placed in the lake with roots; this could also have contributed to the nutrient enrichment of the lake at that time (odgaard 1994). between c. ad 600 and 900 the soil erosion rate to the lake decreased, in agreement with the reduced anthropogenic impact on the landscape as reflected in the decreasing nap values. at this time, the use of the lake for hemp-retting almost ceased and the percentage value of pediastrum sp. declined. between ad 950 and 1050, i.e. in the second half of the viking age, the pollen data suggest renewed deforestation followed by greatly intensified arable farming (elevated percentages of cultivated taxa). this marked change in landuse led to an unprecedented increase in soil erosion that likely caused eutrophication of the lake involving blooms of pedia strum sp. the nutrient enrichment of the lake may have been further intensified by resumption of hemp-retting in the lake, a procedure which continued to around ad 1900. during the medieval period, the anthropogenic impact on the landscape increased significantly as land clearance in 39 2000 mes oli th ic neo lith ic br on ze a ge iro n a ge med iev al mod er n 40 20 sa r -m in ( m g/ cm 2 pe r yr ) 10% 1500% 1000% 500% 10% 4000 3500 3000 2500 2000 1500 1000 500 5000 1000 1500 60 40 20 age (calendar years bc/ad) n a p (% ) sar-min sar-min x 10 nap (%) ∑ pollen from cultivated taxa pediastrum sp. plantago lanceolata a b c d fig. 3. a: sediment accumulation rate of mine ro genic matter (sar-min; proxy of soil erosion rate) and non-arboreal pollen (nap; proxy for open land). the results of the pollen analysis are presented as percentages of the total sum of pollen and spores from terrestrial plants. b: pollen percentages of the sum of cultivated taxa. c: pollen percentages of plantago lanceo lata (ribwort plantain). d: percentages of the green algae pediastrum sp. (calculated on the basis of the terrestrial pollen and spore sum). for b, c and d the dark-coloured curves are percentages and the light-coloured areas represent 10 times exaggeration. rosa_2008:rosa-2008 01/07/09 15:48 side 39 40 creased and crop cultivation intensified. the cultivation of large new areas was facilitated by the introduction of new farming technology such as the mouldboard plough and ridge-and-furrow. at gudme sø, the intensive arable farming in the middle ages led to steadily increasing sediment yields. the abundance of pediastrum sp. at this time suggests continued eutrophication. in modern time (ad 1536 to the present) the area around the lake was further cultivated at the expense of woodland. from c. ad 1500 to 1950, nap values increased from c. 30% to 55% indicating widespread agricultural activities predominantly characterised by crop cultivation. the high and in creasing human impact during this period corresponds closely to a dramatic and unprecedented increase in soil erosion, which peaked around ad 1950 contemporaneously with the highest nap values in the entire record. a rapid and marked increase in the percentage of pediastrum sp. at the beginning of the 18th century signifies the onset of a major nutrient enrichment of the lake. concluding remarks the investigation of the lake sediments from gudme sø provides a 6000-year-record of soil erosion rates as well as contemporary pollen-inferred changes in vegetation and land-use. the study demonstrates a close correlation between changes in agricultural activity and rates of soil erosion in the catchment area of the lake. prior to the late viking age, erosion to the lake occurred in pulses triggered by episodes of land clearance followed by periods of relative stability in the catchment area. after the late viking age, soil erosion accelerated continuously and reached a peak in the mid-20th century with erosion rates approximately 30 times higher than the pre-disturbance rates in the mesolithic. the study dem onstrates that after the introduction of agriculture, soil erosion was mainly caused by human activity rather than cli mate. this general result has also emerged from other lake-based studies of long-term erosion in denmark (ras mussen & bradshaw 2005) and southern sweden (dearing 1991). according to current climate models we can expect an in crease in precipitation, more rainfall in the winter season and a higher frequency of extreme precipitation events in the future (christensen et al. 2006). given these predictions, climate will undoubtedly be of growing significance as a forcing mechanism for soil erosion. acknowledgement the gudme sø investigation is part of an archaeological research project funded by a.p. møller og hustru chastine mc-kinney møllers fond til almene formaal. references behre, k.-e. 1981: the interpretation of anthropogenic indicators in pollen diagrams. pollen et spores 23, 225–245. björck, s. & wohlfarth, b. 2001: 14c chronostratigraphic techniques in paleolimnology. in: last, w.m. & smol, j.p. (eds): tracking environmental change using lake sediments: basin analysis, coring, and chronological techniques 1, 205–245. blaauw, m. & christen, j.a. 2005: radiocarbon peat chronologies and environmental change. journal of the royal statistical society, series c (applied statistics) 54, 805–816. bradshaw, e.g., rasmussen, p., nielsen, h. & anderson, n.j. 2005: midto late-holocene land-use change and lake development at dallund sø, denmark: trends in lake primary production as reflected by algal and macrophyte remains. the holocene 15, 1130–1142. christensen, j.h., christensen, o.b., guldberg, a. & stendel, m. 2006: modeller for klimaets udvikling. in: søndergaard, m., kronvang, b., pejrup, m. & sand-jensen, k. (eds): vand og vejr om 100 år – klima forandringer og det danske vandmiljø, 40–54. højbjerg: forlaget hovedland. dearing, j.a. 1991: erosion and land use. in: berglund, b.e. (ed.): the cultural landscape during 6000 years in southern sweden. ecological bulletins 41, 283–292. european environment agency 2000: down to earth: soil degradation and sustainable development in europe. a challenge for the 21st century. environmental issue series 16, 32 pp. luxembourg: office for official publications of the european communities. gaillard, m.-j. 2007: pollen methods and studies – archaeological applications. in: elias, s.a. (ed.): encyclopedia of quaternary science 3, 2570–2595. oxford: elsevier science. iversen, j. 1973: the development of denmark’s nature since the last glacial. danmarks geologiske undersøgelse v. række 7-c, 126 pp. mackereth, f.j.h. 1966: some chemical observations on post-glacial lake sediments. philosophical transactions of the royal society of london, series b, 250, 165–213. odgaard, b.v. 1993: the sedimentary record of spheroidal carbonaceous fly-ash particles in shallow danish lakes. journal of paleolimnology 8, 171–187. odgaard, b.v. 1994: the holocene vegetation history of northern west jutland, denmark. opera botanica 123, 171 pp. rasmussen, p. & bradshaw, e. 2005: midto late-holocene land-use change and lake development at dallund sø, denmark: study aims, natural and cultural setting, chronology and soil erosion history. the holocene 15, 1105–1115. stone, m. (ed.) 2000: the role of erosion and sediment transport in nutrient and contaminant transfer. international association of hydro logical sciences publication 263, 308 pp. van der knijff, j.m., jones, r.j.a. & montanarella, l. 2000: soil erosion risk assessment in europe, 34 pp. eur 19044 en. ispra: european soil bureau, joint research centre and space applications institute. veihe, a., hasholt, b. & schiøtz, i.g. 2003: soil erosion in denmark: processes and politics. environmental science & policy 6, 37–50. authors’ addresses p.r., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: per@geus.dk j.o., department of earth sciences, aarhus university, høegh-guldbergs gade 2, dk-8000 århus c, denmark. rosa_2008:rosa-2008 01/07/09 15:48 side 40 freiesleben et al. 2024: geus bulletin 57. 8358. https://doi.org/10.34194/geusb.v57.8358 1 of 10 research article beach-ridge formation as a possible indicator for an open limfjord – north sea connection trine h. freiesleben*1,2 , lasse r. berntsen1 , maria blæsbjerg1 , emilia høffer1 , christian f. rasmussen3,4 and nicolaj k. larsen5 1department of natural science (inm), roskilde university, roskilde, denmark; 2department of physics, technical university of denmark, dtu risø campus, roskilde, denmark; 3unesco global geopark, vestjylland, denmark; 4department of ecoscience, aarhus university, aarhus, denmark; 5globe institute, university of copenhagen, copenhagen, denmark abstract raised beach ridges are prograded sequences of wave-built deposits that may provide valuable information about past relative sea-level changes, climate change and coastal evolution. in the limfjord in northern denmark, the early and middle holocene sea-level changes are well-constrained. however, our un derstanding of late holocene sea-level fluctuations is limited, and the exact period when the coastal barrier between the limfjord and the north sea formed remains uncertain. in this study, we use optically stimulated luminescence (osl) dating to determine the age of raised beach ridges at gjellerodde in the western part of the limfjord. the osl ages presented here indicate that the beach ridges formed during three periods at 3.3–2.7, 1.4–1.0, 0.2–0.1 ka. in addition our data suggest a c. 0.2 mm/yr relative sea-level fall during the late holocene. the three distinct peri ods of beach-ridge formation coincide with periods when the limfjord was open towards the north sea as documented in historical records and marine records. this suggests that osl dating of beach ridges can be used as a potential indicator for determining when the connection between the limfjord and the north sea was open in the late holocene. *correspondence: trihof@dtu.dk received: 30 jun 2023 revised: 16 oct 2023 accepted: 30 oct 2023 published: 22 jan 2024 keywords: optically stimulated luminescence, relative sea-level change, beach-ridge formation, limfjord, gjellerodde abbreviations: a.s.l.: above sea level bp: before present (i.e. before 1950) ce: common era dem: digital elevation model ḋ: environmental dose rate ḋe: equivalent dose drc: dose response curve hf: hydrofluoric ir: infrared irsl: infrared stimulated luminescence iqr: inter quartile rejection ka: thousand years osl: optically stimulated luminescence pirir: post-ir irsl pttl: photo-transferred tl rse: relative standard error rsl: relative sea level sar: single-aliquot regenerative dose tl: thermoluminescence w.c.: water content geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: william colgan, geus, denmark reviewed by: lasse sander (alfred wegener institute, germany) and one anonymous reviewer funding: see page 9 competing interests: see page 9 additional files: none 1. introduction humans and habitats in coastal regions are vulnerable to sea-level changes, coastal erosion and climate change that may impact protected fjord environments by opening or closing coastal barriers to the open sea (goodwin et al. 2006; sander et al. 2016). exploring how fjord environments have adapted to natural climate shifts may provide valuable information about how it may evolve in the future as a response to global climate change with rising sea levels and severe storms (shukla et al. 2022). the limfjord in northern denmark is an estuary that is presently connected to the north sea in the west by an artificially maintained channel (thyborøn kanal) dividing the coastal barrier called harboøre tange (south) and agger tange (north) which are collectively referred to as limfjordstangerne (fig. 1). since the last deglaciation, the limfjord has experienced repeated paleoenvironmental changes because of the intricate interplay between glacio-isotatic uplift and eustatic sea-level changes as well as the development and destruction of coastal barriers (bennike et al. 2019). the postglacial sea-level history in the limfjord area is based on radiocarbon dating of terrestrial-to-marine transitions in sediment cores, and mollusc shells in raised marine deposits showing a rapid sea-level rise in the early holocene (petersen 1979; bennike et al. 2019; jessen et al. 2019) followed by a middle holocene peak in sea level 2–5 m higher than present as documented by the raised beach ridges and littorina shorelines (mertz 1924). in the middle holocene during the littorina sea stage (penney 1985), large https://doi.org/10.34194/geusb.v57.8358 https://orcid.org/0000-0002-8616-2099 https://orcid.org/0009-0002-7871-4727 https://orcid.org/0009-0007-9264-2359 https://orcid.org/0009-0003-3724-3719 https://orcid.org/0000-0003-2497-2990 https://orcid.org/0000-0002-0117-1106 mailto:trihof@dtu.dk https://creativecommons.org/licenses/by/4.0/deed.ast https://creativecommons.org/licenses/by/4.0/deed.ast freiesleben et al. 2024: geus bulletin 57. 8358. https://doi.org/10.34194/geusb.v57.8358 2 of 10 geusbulletin.org wide connections existed between the limfjord and the north sea, skagerrak and kattegat (jessen 1910). these connections became shallower or were completely closed by a combination of long-shore sediment transport and con tinued glacio-isostatic uplift that outpaced the eustatic sea-level rise in the late holocene. the closing of the gateways resulted in lower salinity and brackish conditions in the limfjord. however, when the gateways shrank they also became sensitive to the development and destruction of coastal barriers between the limfjord and the north sea. this interplay between sea-level changes and coastal processes led to several periods with fully marine conditions when the natural gateways to the west were open, interrupted by periods of more brackish water when the connec tions were closed (kristensen et al. 1995; lewis et al. 2013). both the relative sea-level (rsl) changes and physical properties of the limfjord are well-constrained in the early and middle holcoene (bennike et al. 2019; jessen et al. 2019; kristiansen et  al. 2021). however, less is known about the late holocene rsl and their influence on the development of the limfjord. in particular, it is uncertain when the limfjordstangerne became established and created a barrier toward the north sea and how stable this barier was in the last 2000–3000 years (kristiansen et  al. 2021). historical documents show that the limfjord was separated from the north sea by a coastal barrier (limfjordstangerne) for an extended time prior  to 1825 ce and it was only interrupted by short periods with open passage (gram-jensen 1991). raised beach ridges constitute a sedimentary archive that can be used to reconstruct holocene sea-level changes, coastal evolution and climate changes (c.f. sander et al. 2016). this method has successfully been used to reconstruct the middle and late holocene in numerous places around the world, including denmark (tamura et al. 2008; nielsen & clemmensen 2009; clemmensen et al. 2012; hede et al. 2015). in this study, we use optically stimulated luminescence (osl) to date raised beach ridges from gjellerodde near lemvig in the western part of limfjorden. our aim is to: (1) constrain the late holocene rsl changes, and (2) link the formation of beach ridges to periods when there was an open connection between the limfjord and the north sea. 2. study site gjellerodde is a cuspate foreland in jylland, west denmark (fig. 1). it is located at the entrance of lemvig, a tributary fjord of the 150 km long limfjord estuary that separates vendsyssel (to the north) from the rest of jylland. gjellerodde lies east of a 40 m high moraine plateau (gjellerbakker) and is a relatively flat low-lying area, primarily covered by heath and small patches of localised forest. a fossil coastline representing the littorina sea cliff is cut into the moraine plateau (jessen 1910; mertz 1924; jessen 1936). gjellerodde consists of postglacial marine sediments composed of numerous gravelly beach ridges. the northern coast is affected by human activity, as the danish coastal authority engages in beach nourishment to protect against erosion. gjellerodde is bisected by an fig. 1 overview of the study area. left: map of gjellerodde in relation to the opening to the north sea at thyborøn kanal. right: aerial photograph of gjellerodde overlain with the digital elevation model (dem) hillshade map. sampling site locations are marked in red, alongside the sample name. the coastal barrier through which the thyborøn kanal flows, consists of harboøre tange and agger tange; collectively known as limfjordstangerne. gjeller sø follup odde gjellerodde thyborøn kanal limfjorden lem vig go-01 go-02 go-03 go-04 go-05 go-06 go-07 go-08 go-09 go-10 go-11 go-12 0 2.5 5 m 0 250 500 km harboøre tange agger tange jylland https://doi.org/10.34194/geusb.v57.8358 freiesleben et al. 2024: geus bulletin 57. 8358. https://doi.org/10.34194/geusb.v57.8358 3 of 10 geusbulletin.org artificial levee, built between 1869 and 1871 (skovborg 1957). trenches were dug into and in between some of the beach ridges during world war ii. the trenches are located between sites go-05 and go-06 and were avoided during sampling. today the area is protected as a bird habitat under the natura 2000 program, administered by lemvig municipality and is used as a grazing area for sheep. the south-eastern section of gjellerodde, known as follup odde, is presently expanding. follup odde is first visible on maps from 1839 (branner 1839) and likely only started forming in the beginning of the 1800s (jessen 1936). from the morphology of gjellerodde and the beach ridges it seems that the present-day sediment transport is controlled by two coastal currents. one current flows west to east from thyborøn canal through limfjorden. the other current within lem vig flows north to south. these two currents form a right angle of sediment transport. the sample sites were chosen to give the best possible indication of the chronological formation of gjellerodde, by choosing sites located on the different ridges. a total of 19 sediment samples were collected from gjellerodde. figure 1 shows an aerial photograph overlain with the digital elevation model (dem) hillshade map to enhance ridge visibility. in fig. 1 the sampling sites are marked in red with the site name. coordinates and sampling depth were noted for each sample location. at sites go-1, go-5 and go-6, multiple samples were taken to test reproducibil ity. samples were taken by digging approximately 60–70 cm vertically into the ridge. to enable infinite dose matrix interpretation of measured dose rates the sample is taken >30 cm from non-homogeneous sediment layers. samples were taken by inserting a metal pipe (length: 20–25 cm, diameter: 5 cm) horizontally into the ridge. the ends were sealed to prevent light contamination to the sample. the sediment samples from go-01 to go-07 consist of  coarse-grained sand and gravel whereas samples go-08 to go-12 consist mainly of coarse-grained sand. 3. methods 3.1 sample preparation the samples were prepared according to standard procedures under subdued red-orange light (aitken 1985). the outer approximately 5 cm of each sample tube was removed and used to estimate the concentration of radionuclides and water content (w.c.). in order to isolate quartz and potassium-rich feldspar from the sediment within each sample tube, a series of steps were carried out (aitken 1985) as follows: (1) the sediment from the inner part of each sample tube was sieved using a mesh size of 180–250 μm. (2) the sieved sediment was subjected to a cleaning process using hydrochloric acid and hydrogen peroxide. (3) to remove the outer 10 μm layer of grains affected by alpha irradiation, a chemical treatment involv ing 10% hydrofluoric (hf) acid was employed. (4) heavy liquid separation was performed (density of 2.58 g/cm3) to separate quartz and potassium feldspar. (5) finally, the extracted quartz was further treated with a 40% hf acid solution. 3.2 measurement facilities the luminescence measurements were conducted with risø thermoluminescence (tl)/osl readers equipped with blue (λ = 470 nm, approximately 80 mw/cm2) and ir (λ = 870 ± 40 nm, approximately 130 mw/cm2) stimulation light sources, as described by bøtter-jensen et al. (2010). for detecting the emitted luminescence emi 9635qa photomultipliers were employed accompanied by 7.5 mm thick hoya u-340 glass filters. the beta irradiation was performed using calibrated 90sr/90y sources mounted on the readers, following the procedures outlined by bøtter-jensen et al. (2010) and hansen et al. (2015). to measure the bulk-radionuclide concentrations, high-resolution gamma spec trometry was employed, as described by murray et al. (1987) and murray et al. (2018). in order to prevent radon loss and ensure a consistent counting setup, sed iment samples were heated to 50°c and embedded in wax, following the method ology outlined by murray et al. (1987). to allow sufficient time for equilibrium between 222rn and its parent 226ra before measurement, the cast samples were stored for a minimum of 3 weeks, as recommended by murray et al. (1987). the conversion of bulk-radionuclide concentrations into infinite matrix dose rates was accomplished by applying the conversion factors and grain size attenuation factors established by guérin et al. (2011) and guérin et al. (2012), respectively. for the internal quartz alpha dose rate, a value of 0.020 ± 0.010 gy/ka was assumed based on vandenberghe et al. (2008). cosmic ray dose rates were cal culated according to prescott and hutton (1994), considering the current burial depths and assuming a 5% uncertainty. the long-term w.c. was set at 18% of the saturated w.c., which was determined as the average value for all samples. an uncertainty of 4% was assumed for the w.c. changing the long-term w.c. value from 0 to 100% of saturated w.c. does not reveal any minimum relative standard error (rse). consequently, when analysing errors, it is not possible to draw any conclusions regarding which assumption is the best. using the current w.c. also does not change the rse. 3.3 optically stimulated luminescence the osl dating method is useful for estab lishing the absolute chronology of most sediments, revealing when they were last exposed to daylight. osl exploits the behaviour https://doi.org/10.34194/geusb.v57.8358 freiesleben et al. 2024: geus bulletin 57. 8358. https://doi.org/10.34194/geusb.v57.8358 4 of 10 geusbulletin.org of electrons within the crystal structures of quartz and feldspar. when sediment is buried, shielding it from sunlight, mineral grains (quartz and feldspar) within the sediment accumulate a latent luminescence signal. this accumulation occurs due to the interaction with ionising radiation emitted primarily by naturally occurring uand th-series, k-40, and cosmic rays. by subjecting the grains to optical stimulation, the in tensity of the osl signal can be measured and calibrated in terms of absorbed dose, known as the equivalent dose (ḋe) in gy. to determine the burial age, the ḋe is divided by the environmental dose rate (ḋ), expressed in gy/ka. the environmental dose rate is determined through independent measurements such as high-resolution gamma spectrometry. the calculation of equivalent doses in this study employed the single-aliquot regenerative dose (sar) procedure, as es tablished by murray and wintle (2000; see section 3.3.1). during the formation of the beach ridge, sediment deposition is assumed to have resulted in exposure to sunlight. this exposure, known as bleaching, acts as a resetting mechanism, providing a measurable age of the sediment. 3.3.1 multi-grain quartz osl measurements dose measurements were conducted on multi-grain quartz aliquots using the sar procedure, as described by murray & wintle (2000). blue light stimulation was applied for 40 s at 125°c to ac quire the measurements, with the intention of avoiding significant accumulation of photo-transferred thermoluminescence (pttl) in the 110°c peak. to minimise potential recu peration effects, a 280°c blue stimulation for 40 s was included between each sar cycle, following the approach outlined by murray & wintle (2003). the aliquots were prepared on stainless steel cups using an 8 mm spot of silicon oil, resulting in approximately 2000 grains per aliquot, as described by duller (2008). to optimise the contribution of the fast-component in the osl signal used for calculations, an early background subtraction technique was employed based on the method proposed by ballarini et al. (2007). the main objective of the sar protocol is to determine the radiation dose equiva lent to the natural luminescence signal measured from the sample. the estimation of equivalent dose relies on a minimum of three sensitivity-corrected regeneration points, along with a recuperation point and a recycling point. individual dose-response curves (drcs) were fitted using a single saturation exponential function that passes through the origin, as described in murray et al. (2021). to derive equivalent doses, interpolation of the sensitivity-corrected natural signal on the individual drcs was performed. average equivalent dose values were calculated using an unweighted arithmetic mean, following the methodology outlined by guérin et al. (2017). the provided uncertainty represents the standard error at a 68% confidence level. to identify and eliminate individual dose values that deviated significantly, the inter-quartile rejection (iqr) criterion, proposed by medialdea et al. (2014), was applied. dose values exceeding 1.5 interquartile ranges above the upper quartile (75%) or below the lower quartile (25%) were rejected. to determine the most suitable sar protocol for this study, a preheat plateau test and a dose-recovery preheat plateau test were conducted. different preheat temperatures ranging from 160 to 260°c, with a cut heat of 40°c, were tested. a stable plateau up to a preheat 220°c was observed in the preheat plateau test. the estimated dose-to-given dose ratio was evaluated with a given dose of c. 6 gy. the selected protocol in this study involved a preheat temperature of 200°c, with a recovery ratio and an infrared (ir) depletion ratio consistent with unity (1.00 ± 0.04, n = 6) and (0.99 ± 0.04, n = 6), respectively. no signifi cant thermal transfer was observed (0.021 ± 0.017 gy, n = 6). additionally, an insignificant recuperation of 0.1 ± 0.3% (n = 6) was observed. 3.3.2 multi-grain feldspar infrared stimulated luminescence (irsl) measurements in all k-rich feldspar dose measurements, a preheat temperature of 250°c was applied for a duration of 100 s. the multi-grain k-rich feldspar aliquots were stimulated using ir light at 50°c for 200 s, followed by an additional ir stimulation at 225°c for 200 s. this particular stimulation protocol, known as ir50 and post-ir irsl 225 (pirir225), was adopted based on the methodology in troduced by buylaert et  al. (2012). to minimise potential recuperation effects, a test dose approximately equal to the natural dose magnitude was employed, and a high-temperature ir bleach was performed at 265°c for 200 s between successive sar cycles. for dose estimation, the light emitted during the initial 2 s of stimulation was considered, while the contribution from the last 10 s of stimulation was subtracted using a late background subtraction (lbg) approach. the aliquots were prepared on stainless steel cups, using a 2 mm spot of silicon oil, resulting in an approximate count of 100 grains per aliquot, following the methodology described by duller (2008). 4. results the radionuclide concentrations derived by using high-resolution gamma spec trometry are summarised in table 1. the dry infinite matrix beta and gamma dose rates are also given, derived assuming a 20 ± 10% loss of 222rn compared to its parent 226ra. total dose rates, also given in table 1, include the contribution from cosmic rays, internal dose rates, and the effects of w.c. 18% of saturated w.c. https://doi.org/10.34194/geusb.v57.8358 freiesleben et al. 2024: geus bulletin 57. 8358. https://doi.org/10.34194/geusb.v57.8358 5 of 10 geusbulletin.org equivalent doses and ages are given in table 2. the ir50 ages were fading corrected using a g-value of 2.82 ± 0.19%/decade (n = 6). figure 2 shows feldspar irsl ages (feldspar) plotted versus osl ages (quartz). the bleaching rate of feldspar is lower than that of quartz. thus the relation between feldspar irsl ages and quartz osl ages can account for the extent of bleaching of the quartz grains at the time of deposition. the feldspar components of the sediment sample should preferably not be older than the quartz. table 1 dose rate summary. sample no. site no. depth (cm)a w.c. (%) radionuclide concentrations (bq/kg) infinite matrix dry dose rates (gy/ka) total dose rates (gy/ka) 238u 226ra 232th 40k beta gamma kfb qb 224501 go-01 41 6 15 ± 2 8.2 ± 0.4 5.9 ± 0.3 220 ± 6 0.68 ± 0.02 0.300 ± 0.007 2.01 ± 0.08 1.08 ± 0.05 224502 go-01 45 5 8 ± 2 6.6 ± 0.2 8.8 ± 0.2 242 ± 5 0.735 ± 0.013 0.340 ± 0.005 2.11 ± 0.08 1.17 ± 0.06 224503 go-01 50 5 7 ± 2 6.2 ± 0.3 6.7 ± 0.2 273 ± 5 0.796 ± 0.013 0.337 ± 0.007 2.15 ± 0.09 1.21 ± 0.06 224504 go-02 45 6 9 ± 3 4.7 ± 0.3 4.8 ± 0.3 355 ± 7 0.97 ± 0.02 0.369 ± 0.008 2.31 ± 0.09 1.38 ± 0.07 224505 go-03 30 6 3 ± 2 5.1 ± 0.4 4.6 ± 0.2 251 ± 6 0.72 ± 0.02 0.288 ± 0.007 2.05 ± 0.08 1.11 ± 0.05 224506 go-04 34 5 6 ± 2 4.9 ± 0.2 4.8 ± 0.2 278 ± 5 0.783 ± 0.014 0.310 ± 0.006 2.12 ± 0.08 1.19 ± 0.06 224507 go-05 37 5 3 ± 2 5.4 ± 0.3 7.7 ± 0.2 247 ± 5 0.729 ± 0.013 0.323 ± 0.007 2.09 ± 0.08 1.15 ± 0.06 224508 go-05 38 6 7 ± 2 8.6 ± 0.2 9.5 ± 0.2 217 ± 4 0.697 ± 0.012 0.343 ± 0.008 2.08 ± 0.08 1.14 ± 0.05 224509 go-05 22 6 5 ± 2 4.7 ± 0.3 5.1 ± 0.2 205 ± 5 0.599 ± 0.013 0.254 ± 0.006 1.93 ± 0.08 1.00 ± 0.05 224510 go-06 44 5 7 ± 2 6.3 ± 0.3 7.7 ± 0.2 297 ± 6 0.865 ± 0.015 0.369 ± 0.008 2.25 ± 0.09 1.31 ± 0.06 224511 go-06 52 5 6 ± 2 6.2 ± 0.3 7.6 ± 0.2 304 ± 5 0.881 ± 0.014 0.372 ± 0.007 2.26 ± 0.09 1.32 ± 0.07 224512 go-06 45 5 8 ± 2 6.3 ± 0.3 8.2 ± 0.3 307 ± 6 0.89 ± 0.02 0.383 ± 0.008 2.28 ± 0.09 1.35 ± 0.07 224513 go-06 75 5 7 ± 2 4.4 ± 0.3 5.5 ± 0.2 258 ± 6 0.73 ± 0.02 0.299 ± 0.007 2.04 ± 0.08 1.11 ± 0.05 224514 go-07 48 5 7 ± 2 5.7 ± 0.3 7.1 ± 0.2 172 ± 5 0.538 ± 0.013 0.259 ± 0.007 1.86 ± 0.08 0.92 ± 0.04 224515 go-08 42 5 5 ± 2 5.1 ± 0.4 5.3 ± 0.2 152 ± 5 0.471 ± 0.014 0.218 ± 0.007 1.77 ± 0.07 0.84 ± 0.04 224516 go-09 40 5 7 ± 2 6.1 ± 0.3 7.3 ± 0.2 216 ± 5 0.654 ± 0.014 0.299 ± 0.007 2.00 ± 0.08 1.06 ± 0.05 224517 go-10 40 5 6 ± 2 6.1 ± 0.3 6.9 ± 0.2 224 ± 5 0.674 ± 0.014 0.301 ± 0.007 2.02 ± 0.08 1.08 ± 0.05 224518 go-11 32 6 8 ± 2 7.1 ± 0.3 7.4 ± 0.3 197 ± 5 0.617 ± 0.014 0.291 ± 0.008 1.97 ± 0.08 1.03 ± 0.05 224519 go-12 30 6 5 ± 3 3.7 ± 0.6 3.9 ± 0.4 124 ± 7 0.38 ± 0.02 0.169 ± 0.009 1.66 ± 0.07 0.72 ± 0.03 dry infinite matrix dose rates are derived from radionuclide concentrations measured using high-resolution gamma spectrometry. total dose rates include the contribution from cosmic rays, internal dose rates and the effects of water content (w.c.). adepths below current surface are given in cm. bq and kf refer to dose rates of quartz and k-rich (12.60 ± 0.15%) feldspar, respectively. table 2 summary of multi-grain quartz results (osl) and multigrain potassium-rich feldspar results (ir50 and pirir225). sample no. site no. elevation (m a.s.l.) dose (gy)a age (ka)b ir50 nr na pirir225 nr na osl nr na ir50 pirir225 osl 224501 go-01 1.56 5.5 ± 0.2 1 11 10.5 ± 0.5 1 11 3.51 ± 0.10 2 16 3.7 ± 0.3 5.0 ± 0.3 3.26 ± 0.19 224502 go-01 1.80 5.23 ± 0.09 2 10 9.7 ± 0.5 1 10 3.79 ± 0.11 1 14 3.5 ± 0.2 4.6 ± 0.3 3.2 ± 0.2 224503 go-01 1.99 6.04 ± 0.19 0 18 10.2 ± 0.4 2 16 4.5 ± 0.3 2 21 3.9 ± 0.2 4.8 ± 0.3 3.7 ± 0.3 224504 go-02 1.62 6.22 ± 0.11 0 12 11.32 ± 0.18 1 11 4.14 ± 0.09 0 18 3.61 ± 0.19 4.9 ± 0.2 3.01 ± 0.17 224505 go-03 1.42 5.38 ± 0.21 0 12 9.1 ± 0.4 0 12 3.86 ± 0.17 1 17 3.32 ± 0.15 4.4 ± 0.3 3.5 ± 0.2 224506 go-04 2.08 5.22 ± 0.09 1 11 9.6 ± 0.3 0 12 3.50 ± 0.11 0 18 3.36 ± 0.18 4.5 ± 0.3 2.95 ± 0.18 224507 go-05 2.92 4.96 ± 0.12 0 12 8.0 ± 0.2 1 11 3.18 ± 0.07 2 16 3.17 ± 0.17 3.8 ± 0.2 2.76 ± 0.15 224508 go-05 3.10 4.66 ± 0.07 3 9 8.3 ± 0.2 0 12 3.01 ± 0.08 0 18 3.05 ± 0.18 4.0 ± 0.2 2.64 ± 0.15 224509 go-05 2.89 4.57 ± 0.15 0 12 7.7 ± 0.5 0 12 2.74 ± 0.07 2 16 3.15 ± 0.16 4.0 ± 0.3 2.75 ± 0.15 224510 go-06 2.68 6.50 ± 0.19 2 10 14.2 ± 1.1 0 12 3.62 ± 0.09 1 17 3.9 ± 0.2 6.3 ± 0.6 2.77 ± 0.16 224511 go-06 2.62 6.13 ± 0.19 0 12 11.3 ± 0.5 0 12 3.70 ± 0.12 1 17 3.64 ± 0.18 5.0 ± 0.3 2.81 ± 0.18 224512 go-06 2.59 7.3 ± 0.3 0 12 12.9 ± 0.8 1 11 3.69 ± 0.08 1 17 4.1 ± 0.3 5.7 ± 0.4 2.74 ± 0.16 224513 go-06 2.59 6.1 ± 0.2 1 11 11.4 ± 0.5 0 12 3.34 ± 0.09 0 18 3.9 ± 0.2 5.6 ± 0.3 3.02 ± 0.18 224514 go-07 2.81 4.67 ± 0.14 0 12 7.6 ± 0.2 2 10 2.55 ± 0.10 0 18 3.42 ± 0.18 4.1 ± 0.2 2.75 ± 0.18 224515 go-08 1.49 2.99 ± 0.17 1 9 6.6 ± 0.7 0 10 2.5 ± 0.4 3 15 2.21 ± 0.15 3.7 ± 0.4 3.0 ± 0.5 224516 go-09 1.44 2.99 ± 0.14 2 8 9 ± 2 1 9 1.34 ± 0.07 1 21 2.04 ± 0.14 4.3 ± 1.0 1.26 ± 0.09 224517 go-10 1.10 2.7 ± 0.2 0 5 5.8 ± 0.8 1 4 1.22 ± 0.13 4 18 1.8 ± 0.2 2.9 ± 0.4 1.13 ± 0.14 224518 go-11 1.05 6.5 ± 1.6 1 4 22 ± 3 0 5 0.27 ± 0.05 2 11 4.5 ± 1.2 11 ± 2 0.27 ± 0.05 224519 go-12 0.96 0.87 ± 0.12 0 8 4.3 ± 0.5 0 8 0.10 ± 0.03 3 24 0.70 ± 0.10 2.6 ± 0.3 0.14 ± 0.04 a.s.l.: above sea level. adose is the arithmetic average equivalent dose after the application of the iqr criterion. all uncertainties are reported at the 68% confidence interval. bage ’ is the equivalent dose divided by the total dose rate (see table 1). na: number of aliquots included in the equivalent dose estimation, that is the total number of measured aliquots less the number of dose estimates rejected by the iqr criterion. nr: number of dose estimates rejected by the iqr criterion. the ir50 ages were fading-corrected using a g-value of 2.82 ± 0.19%/decade. the pirir225 ages were not corrected for fading. https://doi.org/10.34194/geusb.v57.8358 freiesleben et al. 2024: geus bulletin 57. 8358. https://doi.org/10.34194/geusb.v57.8358 6 of 10 geusbulletin.org figure 2 shows a good correlation between osl quartz and pirir225 ages and we conclude that the quartz grains were indeed well-bleached at deposition. note the ir50 signals give older ages than the osl quartz and pirir225 signals. this overestimation may result from the correction for fading of the ir signals in these young samples. nevertheless, the osl quartz ages presented here are considered reliable given their agreement with those obtained by pirir225. 4.1 beach-ridge chronology figure 3 shows osl ages (i.e. burial age) at the sample position. as expected, given the spatial locations of the beach ridges from inner to outer locations and beach progradation from oldest to youngest, there is a general decrease in age from north-west to south-east, with few exceptions. figure 4 shows the osl ages alongside sample elevation. the osl ages of all samples decrease fig. 2 ir50 ages (grey squares) and pirir225 ages (white triangles) from k-rich feldspars (kf) versus osl ages from quartz (q). the ir50 ages were fading-corrected using a g-value of 2.82 ± 0.19%/decade. a 1:1 line is shown (black line) with ± 10% (dotted lines). 0 2 4 6 8 10 12 14 0 1 2 3 4 osl age (q) (ka) ir sl a ge (k f) (k a) ir50 and q pirir225 and q 3.26 ± 0.19 3.2 ± 0.2 3.7 ± 0.3 3.01 ± 0.17 3.5 ± 0.2 2.95 ± 0.18 2.76 ± 0.15 2.64 ± 0.15 2.75 ± 0.15 2.77 ± 0.16 2.81 ± 0.18 2.74 ± 0.16 3.02 ± 0.18 2.75 ± 0.18 3.0 ± 0.0 1.26 ± 0.09 1.13 ± 0.14 0.27 ± 0.05 0.14 ± 0.04 moraine plateau littorina shoreline 500 m n fig. 3 osl ages (ka) and one standard error (including systematic errors). samples are colour coded green, blue and red according to the three phases of beach-ridge formation identified at gjellerodde (see fig. 4 for details of these three events). https://doi.org/10.34194/geusb.v57.8358 freiesleben et al. 2024: geus bulletin 57. 8358. https://doi.org/10.34194/geusb.v57.8358 7 of 10 geusbulletin.org from high to low sample elevations, again with few exceptions. all osl ages were analysed using one-way anova  test, which identified three groups of ages at the 95% significant level. the samples within each group are positioned in a way that corresponds to their geographical location, with the oldest group of samples situated at the highest elevation and the youngest at the lowest elevation (fig. 4). the oldest group is located farthest north-west and the youngest towards the south-east (see fig. 3). based on elevation the oldest  group can be divided into two, in which the most elevated beach ridges (white-edged triangles in fig. 4) yield younger ages than some of morphostratigraphically older ridges located at lower elevations (black-edged triangles in fig.  4).  evaluating deposition rate by including only the lowest beach ridges of the oldest group of samples gives deposition rates of c. 0.2 mm/year (solid black line in fig. 4). 5. discussion 5.1 rsl changes in the late holocene the osl ages (now corrected to years before present (bp), i.e. before 1950) for the raised beach ridges are plotted together with the existing rsl data from limfjorden as recently reviewed by bennike et al. (2019) and jessen et al. (2019; fig. 5a). overall, the osl ages fall on the projected sea-level curve which represents a straight  line from a maximum rsl c. 3 m above sea level (a.s.l.; mertz 1924) approximately 7000 years bp to the present-day sea level. however, some of the osl ages from the highest-elevated beach ridges are younger than some of the morphostratigraphically older ridges located at lower elevations (figs 4 and 5b). this indicates that the rsl remained high at c. 3 m a.s.l. until 3 ka followed by a fairly rapid fall – something that has not been recorded in other records from the area (bennike et al. 2019; kristiansen et al. 2021). instead, we suggest that these beach fig. 4 three distinct groups of osl ages identified for beach-ridge formation in gjellerodde. top: histograms and probability distribution functions for the three groups identified in the lower panel, shown in green, blue and red. lower: osl ages (ka) shown with one standard error (including systematic errors) and their elevation (m). grey shading indicates the three distinct groups of ages defined by the range of data in that group including random and systematic errors at the 95% significant level. groups were identified by a one-way anova test on all samples at the 95% significant level. dark green triangles with black outline: samples <2.5 m elevation. green triangles with white outline: samples >2.5 m elevation. see also fig. 3. the solid black line represents a linear interpolation excluding the highest beach ridges. 012345 age (ka) 0.0 0.5 1.0 1.5 2.0 p ro ba bi lit y ×10–3 e le va tio n (m a .s .l. ) 0.0 0.5 1.0 1.5 2.0 2.5 3.0 3.5 https://doi.org/10.34194/geusb.v57.8358 freiesleben et al. 2024: geus bulletin 57. 8358. https://doi.org/10.34194/geusb.v57.8358 8 of 10 geusbulletin.org ridges are unrelated to a specific sea level and formed under an extraordinary storm event that produced some extraordinarily high beach ridges (clemmensen et al. 2012). accordingly, we place the rsl curve through the oldest group of osl ages c. 1.5–2 m a.s.l. (figs 4 and 5a). using a simple linear interpolation that excludes the highest beach ridges suggests a drop in rsl of c. 1 m in the last 3000 years, that is c. 0.2 mm/yr (fig. 4). this estimate is lower than estimates from nørholm in the northern part of the limfjord where the rate of sealevel change was estimated to be c. 1.2 ± 0.2 mm/yr in the last 3000 years (kristiansen et al. 2021). the difference between the two sites is, however, expected as gjelleroddde is located c. 75 km south-west where the average long-term isostatic rates of uplift are lower (c. 0.25 mm/yr) and in line with our estimates (hansen et al. 2012). although our results overall are consistent with previous investigations, we acknowledge that beachridge formation is not only controlled by changes in rsl and that sediment supply and wind regime may also play an important role. 5.2 beach-ridge formation as a potential indicator for an open limfjord – north sea connection the new osl ages presented here indicate that beach ridges did not develop continuously in the late holocene but were formed during three discrete periods 3.3–2.7, 1.4–1.0, 0.2–0.1 ka (fig. 4). however, because only a subset of the beach ridges were dated we cannot rule out the possibility that they also formed in other time periods or that the existing periods should be expanded. the formation of the most recent beach ridges is related to the recent opening in 1825 ce (gram-jensen 1991) and this formation of new beach ridges has continued until the present-day. the second and third succession of beach ridges comprise most of the beach ridges at gjelleroddde and they were formed 3.3–2.7 and 1.4–1.0 ka. both periods coincide with increased salinity in the limfjord indicating full marine conditions and open connections to the north sea (kristensen et al. 1995; lewis et al. 2013). it is interesting to see that some of the beach ridges date to the viking age (793–1066 ce, see the central group (blue) fig. 5 comparison of osl ages from this study with previously published radiocarbon histories and palaeo-salinity data. a: holocene rsl curve (dashed curve) for limfjorden based on radiocarbon-dated samples (bennike et al. 2019; jessen et al. 2019) and osl ages (this study) vs. elevation (m). all ages are in years bp (i.e. before 1950). error bars are one standard error. marine and terrestrial 14c ages are shown in blue and green, respectively. b: the same osl ages are shown alongside palaeo-salinity data for the limfjord (kristensen et al. 1995; lewis et al. 2013). high salinity indicates an open connection towards the north sea (kristensen et al. 1995; lewis et al. 2013). dark grey shading indicates periods when the connection between the limfjord and the north sea was open according to the osl ages (this study) and the light grey shading indicates open periods according to salinity concentrations (kristensen et al. 1995; lewis et al. 2013). 020004000600080001000012000 age (years bp) –45 –40 –35 –30 –25 –20 –15 –10 –5 0 5 e le va tio n (m a .s .l. ) relative sea-level change (a) marine (14c) terrestrial (14c) terrestrial (osl) 010002000300040005000 age (years bp) 0 1 2 3 e le va tio n (m a .s .l. ) (b) 15 25 5 35 s al in ity (‰ ) 30 20 40 10s al in ity (g /l ) cl os ed -o pe n cl os ed -o pe n https://doi.org/10.34194/geusb.v57.8358 freiesleben et al. 2024: geus bulletin 57. 8358. https://doi.org/10.34194/geusb.v57.8358 9 of 10 geusbulletin.org in fig. 4) when it has been speculated that an opening to the west existed (eriksen et al. 2009). according to these accounts, the opening to the west closed shortly after the viking age (eriksen et al. 2009) in agreement with the oslbeach-ridge chronology (1.4–1.0 ka). overall, our data indicate that osl dating of beach ridges may indicate an open limford – north sea connection. the open connection would have facilitated a stronger current that would increase erosion and this would have been the dominant factor in producing the beach ridges at gjellerodde and possibly in a more extensive area in the western limfjord area. other factors may also have influenced the formation of beach ridges including changes in storm intensity, currents, and sediment availability. however, to test this hypothesis would require investigation of more beach ridge systems to link their timing of formation with geomorphological processes in the western limfjord. 6. conclusions in this study, we used osl dating of beach ridges in combination with existing sea-level proxy data to determine late holocene sea-level changes and identify when the limfjord was last connected to the north sea. based on 19 osl ages from beach ridges at gjellerodde, the late holocene rate of rsl change was estimated to be c. 0.2 mm/yr, consistent with other estimates from northern denmark. our osl ages furthermore indicate that the beach ridges at gjellerodde were most likely formed during three discrete time periods at 3.3–2.7, 1.4–1.0, 0.2–0.1 ka. overall there is a good correlation between known historical accounts and paleoenvironmental reconstructions of salinity changes and the formation of beach ridges in the western limfjord. accordingly, this suggests that osl dating of beach ridges can be used as an indicator for determining when there was an open connection between the western limfjord and the north sea in the late holocene. acknowledgements the authors thank lemvig municipality for access to gjellerodde and the danish coastal authority for valuable discussions. they also thank lasse sander and an anonymous reviewer for constructive comments that helped improve the manuscript. additional information author contributions thf, cfr & nkl: conceptualisation of the study and supervision. lrb, mb, eh: fieldwork, osl measurements and data analysis. thf, lrb, mb, eh, cfr & nkl: data interpretation and writing the first draft of the manuscript. competing interests the authors declare that there are no competing interests. funding statement measurement facilities 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3. methods 3.1 sample preparation 3.2 measurement facilities 3.3 optically stimulated luminescence 3.3.1. multi-grain quartz osl measurements 3.3.2 multi-grain feldspar infrared stimulated luminescence (irsl) measurements 4. results 4.1 beach ridge chronology 5. discussion 5.1 relative sea-level changes in the late holocene 5.2 beach-ridge formation as a potential indicator for an open limfjord – north sea connection 6. conclusions acknowledgements additional information references figures fig. 1 overview of the study area. left: map of gjellerodde in relation to the opening to the north sea at thyborøn kanal. right: aerial photograph of gjellerodde overlain with the digital elevation model (dem) hillshade map. sampling site locations are marked in red, alongside the sample name. the coastal barrier through which the thyborøn kanal flows, consists of harboøre tange and agger tange; collectively known as limfjordstangerne. fig. 2 ir50 ages (grey squares) and pirir225 ages (white triangles) from k-rich feldspars (kf) versus osl ages from quartz (q). the ir50 ages were... fig. 3 osl ages (ka) and one standard error (including systematic errors). samples are colour coded green, blue and red according to the three phases of beach-ridge formation identified at gjellerodde (see fig. 4 for details of these three events). fig. 4 three distinct groups of osl ages identified for beach-ridge formation in gjellerodde. top: histograms and probability distribution functions for the three groups identified in the lower panel, shown in green, blue and red. lower: osl ages (ka) shown with one standard error (including systematic errors) and their elevation (m). grey shading indicates the three distinct groups of ages defined by the range of data in that group including random and systematic errors at the 95% significant level. groups were identified by a one-way anova test on all samples at the 95% significant level. dark green triangles with black outline: samples <2.5 m elevation. green triangles with white outline: samples >2.5 m elevation. see... fig. 5 comparison of osl ages from this study with previously published radiocarbon histories and palaeo-salinity data. a: holocene rsl curve (dashed curve) for limfjorden based on radiocarbon-dated samples (bennike et al. 2019; jessen et al. 2019) and osl ages (this study) vs. elevation (m). all ages are in years bp (i.e. before 1950). error bars are one standard error. marine and terrestrial 14c ages are shown in blue and green, respectively. b: the same osl ages are shown alongside palaeo-salinity data for the limfjord (kristensen et al. 1995; lewis et al. 2013). high salinity indicates an open connection towards the north sea (kristensen et al. 1995; lewis et al. 2013). dark grey shading indicates periods when the connection between... tables table 1 dose rate summary. table 2 summary of multi-grain quartz results (osl) and multigrain potassium-rich feldspar results (ir50 and pirir225). research article | short andersen et al. 2023: geus bulletin 53. 8335. https://doi.org/10.34194/geusb.v53.8335 1 of 6 scaling the danish national water resources model for a paneuropean quasi-3d groundwater resources model lærke therese andersen*1  , anne-sophie høyer1  , mette hilleke mortensen1  , lars troldborg2  , klaus hinsby2  1department of near surface land and marine geology, geological survey of denmark and greenland (geus), copenhagen, denmark. 2department of hydrology, geological survey of denmark and greenland (geus), copenhagen, denmark abstract in this study, we upscale and simplify hydrostratigraphic information from a detailed model for denmark to a pan-european scale. this is part of a larger project to develop a harmonised overview of the volume and depth of groundwater resources in a quasi-3d european groundwater resource model. a 10 km grid and a maximum of c. 10 hydrostratigraphic layers were chosen as the common scale for the european database. the danish information is based on the national water resources model (the dk-model), where the information is significantly more detailed (100 m grid and up to 26 layers). information was transferred from the dk-model to the quasi-3d model by a method involving computations of mean volumes and expert assessment to reduce layers in each cell. in this process, detailed hydrostratigraphic information is lost, which could otherwise be used for local groundwater flow modelling in denmark. however, the strength of the quasi-3d model is that it still contains the volumes of all hydrostratigraphic units, both the saturated and unsaturated parts. hence, the upscaled model can contribute to a relatively precise calculation of european groundwater resources for the quantitative assessment of groundwater status across europe at a 10 × 10 km scale. *correspondence: lta@geus.dk received: 23 nov 2022 revised: 14 july 2023 accepted: 3 aug 2023 published: 16 oct 2023 keywords: aquifers, europe, groundwater resources, quasi-3d groundwater resource model, upscaling hydrostratigraphic layer model abbreviations: dk-model: danish national water resources model inspire: infrastructure for spatial information in europe resource: pan-eu groundwater resources map geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: julian koch (geus, denmark) reviewed by: ioannis trichakis (european commission joint research centre, italy) and one anonymous reviewer funding: see page 5 competing interests: none additional files: none introduction in a changing climate, there is an increasing need for a 3d pan-european harmonised overview of european groundwater resources that includes supplementary information on the depth and volume of exploitable groundwater resources, as well as water balance and groundwater mean residence time of the aquifers. the aim of this study was, therefore, to upscale the danish national water resources model (referred to as the dk-model) to contribute to a pan-european quasi-3d groundwater resources model developed as part of the european project resource (pan-eu groundwater resources map). the dk-model is significantly more detailed than the required input for the european model. hence, the information in the dk-model had to be simplified and upscaled. upscaling models is a challenging task, which has been managed in different ways by the many researchers dealing with this task (zhang et al. 2021). these studies have typically been conducted to simulate flow and transport processes in oil or groundwater studies to reduce computational time and power. however, in this first approach of upscaling local models to create a pan-european model, the focus is only on groundwater resources and not on flow or transport processes. in this study, hydrogeological information was transferred from a 100 m grid to a 10 km grid, and from 26 to 13 layers, using a mean-volume approach. the results are discussed and exemplified by one of the 10 × 10 km cells in denmark. https://doi.org/10.34194/geusb.v53.8335 https://orcid.org/0000-0002-4296-2510 https://orcid.org/0000-0002-9105-5554 https://orcid.org/0000-0001-8585-5881 https://orcid.org/0000-0002-7366-1438 https://orcid.org/0000-0003-1190-4550 mailto:lta@geus.dk https://creativecommons.org/licenses/by/4.0/ andersen et al. 2023: geus bulletin 53. 8335. https://doi.org/10.34194/geusb.v53.8335 2 of 6 www.geusbul let in.org the resource (pan-eu groundwater resources map) project was part of the european programme geoera, in which hydrogeologists and groundwater scientists from more than 30 european geological surveys developed a pan-european harmonised overview of the volume and depth of groundwater resources (hollis et al. 2022). two examples of the harmonised pan-european groundwater resources map (kivits et al. 2021) are shown in figs 1a and b. the examples show the aquifer types and the total depths of the freshwater systems in the contributing areas of europe (egdi 2022). both parameters show great variation not only across europe but also within denmark. geological setting the pre-quaternary surface in denmark dips stratigraphically towards the north-east, such that the pre-quaternary is composed of the youngest sediments (pliocene, miocene and palaeogene) in the south-west and oldest (danian and upper cretaceous limestone) in the north-east. the aquifer type (fig. 1a) and the total depth of aquifers across denmark (fig. 1b) are closely related to the setting of the pre-quaternary layers. overlaying these layers, the quaternary deposits are mainly composed of clayey and sandy tills, interglacial deposits and meltwater sand and clay deposits. the quaternary sequence is generally highly complex due to glacial tectonism (e.g. høyer et al. 2013) and buried glacial valleys (sandersen & jørgensen 2017). danish drinking water supply is almost entirely sourced from groundwater and, due to the geology, is extracted from different types of deposits across denmark. in the south-west, groundwater is primarily extracted from miocene sand deposits (highly productive intergranular aquifers; fig. 1a); in central areas, it is primarily from pleistocene meltwater sand deposited in buried valleys; and in the east and north, it is primarily from chalk and limestone of maastrichtian and danian age, respectively. the hydraulic characteristics of the upper part of the chalk and limestone are strongly affected by fissures, brecciation and other features caused by glacial overpressure and glaciotectonics (pedersen et al. 2018), which increase the permeability and hence make them important carbonate aquifers (highly productive fissured aquifers; fig. 1a). the dk-model the main information for the map of the european groundwater resources was derived from the dk-model (stisen et al. 2019). the dk-model is based on four individual models for jylland (jutland), fyn (funen), sjælland (zealand) and bornholm. together, these models represent the aquifers and aquitards in the majority of denmark. the dk-model has been developed to provide a nationwide overview of groundwater resources, defined on a grid of 100 × 100 m. the number of layers vary according to hydrogeology, such that most layers are present in jylland (26 layers; see fig. 2). the number of pre-quaternary layers on top of the limestone increases from the north, where limestone is present close to the surface, and towards the south-west, where limestone is situated at great depth. the model encompasses up to six sandy aquifers in the pre-quaternary and six sandy aquifers in the quaternary deposits. some of the layers appear regionally extensive, whereas others, such as fig. 1 examples of mapped features from the pan-european groundwater resources map (egdi 2022). (a) aquifer types. (b) total depth of aquifers below the surface. the blank areas (grey shading) reflect the countries or regions that did not participate or finalise the grid. denmark is situated within the black square. a) b)a bminor aquifers, essentially no groundwater minor aquifers, little groundwater locally productive intergranular aquifers highly productive intergranular aquifers highly productive fissured aquifers locally productive fissured aquifers 500 km 500 km 100–200 m 200–400 m 50–100 m 400–800 m >800 m <50 m https://doi.org/10.34194/geusb.v53.8335 http://www.geusbulletin.org andersen et al. 2023: geus bulletin 53. 8335. https://doi.org/10.34194/geusb.v53.8335 3 of 6 www.geusbul let in.org the uppermost pre-quaternary clay layer (pc1, fig. 2), only have a local appearance. in the quaternary layers, valley deposits (qc7, qs6, qc6 and qs5, fig. 2) primarily occur locally in individual valleys, whereas layers in the plateaus are regionally extensive. each of the layers in the dk-model is associated with distinct lithology types. the dk-model is used for dynamic simulation of all major parts of the hydrological system, using a mike she/mike1d engine for groundwater calculations. the estimated hydrological parameters for each of the hydrostratigraphic units are optimised during the calibration of the model (stisen et al. 2019). methods our goal was to upscale and simplify the 100 × 100 m dk-model with 26 layers to a 10 × 10 km quasi-3d groundwater resources model with 10 layers without losing information. for example, important aquifers in denmark, such as the buried quaternary valleys, are elongate erosional features (much less than 10 km wide) that cut the prevailing stratigraphy filled with younger sediments (sandersen & jørgensen 2017). the buried valleys are challenging to recreate when upscaling the model to a grid wider than the extent of the valleys, hence a special method is developed (see figs 3 and 4). the quasi-3d model is based on an inspire (infrastructure for spatial information in europe) grid from the european environment agency, consisting of cells each with a unique cell id. several geological and hydrological parameters were applied to each cell, such as total depth of active layers, boundary definition of the active depth and surface level altitude. other layer information included thickness and extent of the unsaturated and saturated zones, hydrogeofacies, geological age, type of layer (aquitard or aquifer) and indications of the presence of palaeo-, artesianor thermal groundwater. hydrological parameters included porosity and horizontal and vertical conductivities for a maximum of c. 10 layers (kivits et al. 2020). simplification of the information in the dk-model was initially performed in jylland, by grouping near-surface quaternary layers (qs1, qc2, qs2; fig. 2) and deeper quaternary layers (qs5, qc6, qs6; fig. 2). this was based on experience with modelling of groundwater mapping in denmark where layers of similar hydraulic properties are grouped together in the dk-model. for both sjælland and jylland, the lower most quaternary till unit (qc7; fig. 2) was grouped together with the top pre-quaternary clay unit (pc1; fig. 2). to further reduce the number of layers, a threshold for the mean thickness (3 m) and the mean extent (5%) of each layer in each cell was defined, as this was considered 0 40 80 120 160 200 0 20 40 60 80 100 extent (%) 0 2000 4000 6000 8000 80 40 0 –40 –80 –120 el ev at io n (m ) distance (m) d ep th (m )5: till & meltwater clay 7: till & meltwater clay 3: till & meltwater clay 1: till 4: sand 2: sand 6: sand 9: limestone 8: paleogene clay a b 1 2 3 4 5 6 7 8 9 fig. 3 example results from a representative cell (cell no. 433_358, for location see fig. 4). (a) representative 2d profile through the dk-model in the cell. (b) extent and calculated mean thicknesses of the saturated and unsaturated part (black bars) of each layer in the cell – here seen in 2d but representing the entire 3d cell (10 × 10 km). fig. 2 principal sketches for the layers in the dk-model for jylland, modified from stisen et al. 2019. qs: quaternary sand. qc: quaternary clay. ps: pre-quaternary sand. pc: pre-quaternary clay. top 2 m distributed after soil map qc1 qs1 qs2 qs3 qs4 qs5 qs6 qc2 qc3 qc4 qc5 qc6 qc7 ps1 ps2 ps3 ps4 ps5 ps6 pc7 pc6 pc5 pc4 pc3 pc2 pc1 sw ne limestone limestone and chalk quaternary clay quaternary sand pre-quaternary sand pre-quaternary clay limestone https://doi.org/10.34194/geusb.v53.8335 http://www.geusbulletin.org andersen et al. 2023: geus bulletin 53. 8335. https://doi.org/10.34194/geusb.v53.8335 4 of 6 www.geusbul let in.org to be the minimum volume of interest at the european scale. even after this reduction, 190 cells still had more than 10 layers. it was therefore decided to allow for up to 13 layers to reduce the number of cells that should be handled manually. layers in the remaining 88 cells that had more than 13 layers were manually reduced by expert judgement. where possible, layers were grouped to ensure only a limited impact on the final resource evaluation, for example, in areas where two aquifers or aquitards were situated directly on top of each other. hydrological parameters were defined based on optimised parameters from the calibrated dk-model (stisen et al. 2019). however, the hydrological parameters had to be adjusted to the new layer grouping. generally, most grouping was performed between aquifers and aquitards, respectively. in these cases, the hydraulic parameters were calculated as a mean value if the parameters were not identical. the other type of grouping was performed in areas with thin layers which have been neglected in the final model result. this is the case, for example, where thick miocene clay layers surround thin miocene sand layers. in these cases, the hydraulic conductivity of the sand layer has not been considered because the final conductivity was dominated by the thick clay layers. next, a python script in arcgis was programmed to extract the mean thickness and extent from the layers in the dk-model to the cells in the quasi-3d model. the unsaturated and, hence, the saturated parts were calculated using 500 × 500 m simulations of depth to the phreatic surface from the dk-model. the total depth of the groundwater system in denmark was defined as the depth to a hydrogeological base, a layer with very low vertical hydraulic conductivity. the top of the hydrogeological base in the western part of denmark was defined as the palaeogene clay, and in the northern and eastern parts was defined as the depth to the compact (non-fissured) cretaceous limestone, estimated to be present at depths below 50 m of the top of the limestone (fig. 1b). metadata and methods used for upscaling the dk-model in the rest of denmark are described by andersen et al. (2021). results the results of the conversion of the dk-model from a hydrostratigraphic layer model to a quasi-3d groundwater resources column model can be seen in fig. 3. a profile crossing one of the 10 × 10 km grid cells on the island of fyn shows nine layers in the dk-model numbered in chronostratigraphic order (fig. 3a). in fig. 3b, the same cell is shown in the quasi-3d model. here, the calculated mean thickness and extents of each layer are also arranged in chronostratigraphic order. the unsaturated parts of the layers are shown with black bars on top of the layers – only the three topmost layers are a b <5 m 5–10 m 10–15 m 15–20 m 20–25 m >25 m mean thickness profile in fig. 3 cell no. 433_358 mean extent profile in fig. 3 cell no. 433_358 <1% 1–5% 5–15% 15–30% 30–45% 45–75% 200 km200 km 10 km 10 km fig. 4 calculated mean thicknesses (a) and mean extents (b) for layer 6 (corresponding to valley sand) in 10 × 10 km cells covering the island of fyn. overlaying this are the exact position and thickness of the valley sand layer extracted from the dk-model, shown as a coloured contour map (colour scale in a). the position of the profile in fig. 3a and the cell in fig. 3b are marked on the map. https://doi.org/10.34194/geusb.v53.8335 http://www.geusbulletin.org andersen et al. 2023: geus bulletin 53. 8335. https://doi.org/10.34194/geusb.v53.8335 5 of 6 www.geusbul let in.org partly unsaturated. in the final quasi-3d model, layers 7 and 8, which are both clay layers, have been joined. note that layers in fig. 3b represent the entire 3d cell. for instance, layer 4 occurs in the entire profile length (10 km; fig. 3a), but not in the entire cell (10 × 10 km; fig. 3b) where the extent is less than 100%. naturally, the mean thicknesses are typically most representative for layers of great extent and little thickness variation (e.g. layer 4) compared to layers of little extent and great thickness variation (e.g. layer 6, a sand layer in a buried glacial valley with a mean thickness less than 10 m; fig. 3b). in practice, though, the thickness of the valley sand varies up to 40 m along the profile (fig. 3a, approx. distance 3000 m). however, the strength of the pan-european quasi-3d groundwater resources model is that the volume of all layers is equally well represented, no matter how extensive the layers are in the 10 × 10 km cell. in fig. 4 we zoom out and see the quasi-3d model for the entire island of fyn with the mean thickness (fig. 4a) and extent (fig. 4b) of layer 6. the two maps are overlaid by the actual occurrence and thickness of the valley sand layer (layer 6 in fig. 3a). the figure illustrates how the local occurrence and thickness of a layer in the dk-model is extrapolated to the entire cell in the quasi-3d model. for example, layer 6 in the dk-model has a local maximum (>25 m, red) where the profile crosses the cell (fig. 3a). however, in the entire cell in the quasi-3d model, it is green, as the thickness when extrapolated to the entire cell is only 5–10 m. because the layer numbering is individual for each 3d cell (kivits et al. 2020), the stratigraphy of layer 6 might vary within the map. however, the stratigraphy is relatively uniform at fyn, and layer 6 therefore almost corresponds to the valley sand in the entire area. discussion the upscaled hydrostratigraphy has been constructed to gather an overview of the groundwater resources on a pan-european scale. but simplifying the hydrogeology from a 100 m grid to a 10 km grid does have implications on the use for hydrological simulations, as the flow resolution from the dk-model is likely to be sensitive to hydrogeological structures at scales less than 10 km. because the layer numbering is individual in each cell (kivits et al. 2020), the quasi-3d model cannot be used directly for groundwater flow and recharge across all cell boundaries. the method of numbering layers is, however, important for all volumes to be represented in the model and placed in the right chronostratigraphic order, regardless of different geological regions and structures. the saturated parts of the aquifer volumes were used in the calculation of the groundwater resource (schoonderwoerd et al. 2021). the most challenging part of the upscaling has been to reduce the number of layers from 26 to 13 in each cell. the decision of 13 layers was based on our definition of ‘the minimum volume of interest on a european scale’ – a definition that will be more consolidated when the pan-european model has actually been used for decision-making in the eu. in the process of reducing the layers, expert judgement played an important role. as this was manually assessed, it was the most time-consuming part of the process, which could have been reduced by increasing the minimum volume. the quasi-3d model cannot be used for detailed groundwater flow simulations. however, simulations of water balance and groundwater mean residence time and assessments of water stresses and climate change impacts are likely to make reasonable sense at a 10 km scale. conclusions we have upscaled and simplified the hydrostratigraphy of the dk-model from the current 100 m grid to 10 km grid for use in a pan-european quasi-3d model. if we had used a traditional hydrostratigraphic model, all details about geological structures smaller than 10 × 10 km would have been lost. however, the upscaled model used here is a groundwater resources column model, showing the extent and thickness of each layer in the cell. therefore, the volumes of features of a small extent, but large thickness, for example, the buried quaternary valleys in denmark, are still represented. further, since the structure of the model has changed from a hydrostratigraphic layer model to a quasi-3d resource model, the upscaling most likely has implications for the types of detailed groundwater flow simulations that the dk-model was originally used for, for example, in groundwater models when analysing contaminant transport or groundwater travel time distributions. as the quasi-3d model considers both the saturated and unsaturated parts of the aquifers, it contributes to a relatively precise calculation of the european groundwater resources and the quantitative status across europe in a 10 × 10 km grid. acknowledgements we would like to thank tano kivits for reading and commenting on the manuscript. we also thank the reviewers for being thorough and giving constructive remarks to the manuscript. the work was co-funded by the european union’s horizon 2020 research and innovation programme under grant agreement no. 731166 and innovation fund denmark under agreement no. 8055-00073b. funding statement the work was co-funded by the european union’s horizon 2020 research and innovation programme under grant agreement no. 731166 and innovation fund denmark under agreement no. 8055-00073b. https://doi.org/10.34194/geusb.v53.8335 http://www.geusbulletin.org andersen et al. 2023: geus bulletin 53. 8335. https://doi.org/10.34194/geusb.v53.8335 6 of 6 www.geusbul let in.org author contributions lta: conceptualisation, formal analysis, methodology, resources, validation, visualisation. ash: writing original draft, visualisation. mhm: formal analysis, methodology, resources. lt: conceptualisation, formal analysis, methodology, resources. kh: funding acquisition, conceptualisation, project administration, supervision, validation. competing interests none additional files none references andersen, l.t., mortensen, m.h., troldborg, l. & hinsby, k. 2021: geoera resource groundwater deliverable 6.6. justification of the choices to compile the pan-eu groundwater resources maps by all participating countries – metadata denmark. 52–62. https://repository. europe-geology.eu/egdidocs/resource/geoera%2bresource%2bdeliverable%2b66%2bjustification%2bof%2bc.pdf (accessed september 2023) egdi (european geological data infrastructure) 2022: pan-european groundwater resource map: volumes and depths of prime european aquifers. https://data.geus.dk/egdi/?mapname=egdi_new_structure#baslay=basemapgeus&extent=-710780,1211800,6797980,52047 10&layers=resource_netcdf11 (accessed november 2022) hollis, j. et al. 2022: pan-european geological data, information, and knowledge for a resilient, sustainable, and collaborative future. european geologist 53, 5–19. https://doi.org/10.5281/zenodo.6883282 høyer, a.s., jørgensen, f., piotrowski, j.a. & jakobsen, p.r. 2013: deeply rooted glaciotectonism in western denmark: geological composition, structural characteristics and the origin of varde hill-island. journal of quaternary science 28(7), 683–696. https://doi. org/10.1002/jqs.2667 kivits, t., janza, m. & broers, h.p. 2020: geoera resource groundwater deliverable 6.2. database with information on volumes and depths at 10x10 and/or 25x25 km grids, 28 pp. https://repository.europe-geology.eu/egdidocs/resource/geoera%2bresource%2bdeliverable%2b62 %2bdatabase%2bwith%2binfo.pdf (accessed september 2023) kivits, t., rasmussen, v. & broers, h.p. 2021: geoera resource groundwater deliverable 6.4. dataset to be included in the information platform, 10 pp. https://repository.europe-geology.eu/ egdidocs/resource/geoera%2bresource%2bdeliverable%2b64%2bdataset%2bto%2bbe%2bincl.pdf (accessed september 2023) pedersen, s.a.s., gravesen, p. & hinsby, k. 2018: chalk-glacitectonite, an important lithology in former glaciated terrains covering chalk and limestone bedrock. geological survey of denmark and greenland bulletin 41, 21–24. https://doi.org/10.34194/geusb. v41.4333 sandersen, p.b.e. & jørgensen, f. 2017: buried tunnel valleys in denmark and their impact on the geological architecture of the subsurface. geus bulletin 38, 13–16. https://doi.org/10.34194/geusb.v38.4388 schoonderwoerd, e., essink, o.g., bos-burgering, l., kivits, t., zaadnoordijk, w.j. & broers, h.p. 2021: geoera resource groundwater – deliverable 6.5. water balance terms for the eu fresh groundwater grid. 32 pp. https://repository.europe-geology.eu/egdidocs/resource/geoera%2bresource%2bdeliverable%2b65%2bwater%2bbalance%2bterm. pdf (accessed september 2023) stisen, s., ondracek, m., troldborg, l., schneider, r.j.m. & john van til, m. 2019: national vandressource model – modelopstilling og kalibrering af dk-model 2019. danmarks og grønlands geologiske undersøgelse rapport 2019/31, 1–125. https://doi.org/10.22008/gpub/32631 zhang, x. et al. 2021: application of upscaling methods for fluid flow and mass transport in multi-scale heterogeneous media: a critical review. applied energy 303, 117603. https://doi.org/https://doi.org/10.1016/j. apenergy.2021.117603 https://doi.org/10.34194/geusb.v53.8335 http://www.geusbulletin.org https://repository.europe-geology.eu/egdidocs/resource/geoera%2bresource%2bdeliverable%2b66%2bjustification%2bof%2bc.pdf https://repository.europe-geology.eu/egdidocs/resource/geoera%2bresource%2bdeliverable%2b66%2bjustification%2bof%2bc.pdf https://repository.europe-geology.eu/egdidocs/resource/geoera%2bresource%2bdeliverable%2b66%2bjustification%2bof%2bc.pdf https://data.geus.dk/egdi/?mapname=egdi_new_structure#baslay=basemapgeus&extent=-710780,1211800,6797980,5204710&layers=resource_netcdf11 https://data.geus.dk/egdi/?mapname=egdi_new_structure#baslay=basemapgeus&extent=-710780,1211800,6797980,5204710&layers=resource_netcdf11 https://data.geus.dk/egdi/?mapname=egdi_new_structure#baslay=basemapgeus&extent=-710780,1211800,6797980,5204710&layers=resource_netcdf11 https://doi.org/10.5281/zenodo.6883282 https://doi.org/10.1002/jqs.2667 https://doi.org/10.1002/jqs.2667 https://repository.europe-geology.eu/egdidocs/resource/geoera%2bresource%2bdeliverable%2b62%2bdatabase%2bwith%2binfo.pdf https://repository.europe-geology.eu/egdidocs/resource/geoera%2bresource%2bdeliverable%2b62%2bdatabase%2bwith%2binfo.pdf https://repository.europe-geology.eu/egdidocs/resource/geoera%2bresource%2bdeliverable%2b62%2bdatabase%2bwith%2binfo.pdf https://repository.europe-geology.eu/egdidocs/resource/geoera%2bresource%2bdeliverable%2b64%2bdataset%2bto%2bbe%2bincl.pdf https://repository.europe-geology.eu/egdidocs/resource/geoera%2bresource%2bdeliverable%2b64%2bdataset%2bto%2bbe%2bincl.pdf https://repository.europe-geology.eu/egdidocs/resource/geoera%2bresource%2bdeliverable%2b64%2bdataset%2bto%2bbe%2bincl.pdf https://doi.org/10.34194/geusb.v41.4333 https://doi.org/10.34194/geusb.v41.4333 https://doi.org/10.34194/geusb.v38.4388 https://repository.europe-geology.eu/egdidocs/resource/geoera%2bresource%2bdeliverable%2b65%2bwater%2bbalance%2bterm.pdf https://repository.europe-geology.eu/egdidocs/resource/geoera%2bresource%2bdeliverable%2b65%2bwater%2bbalance%2bterm.pdf https://repository.europe-geology.eu/egdidocs/resource/geoera%2bresource%2bdeliverable%2b65%2bwater%2bbalance%2bterm.pdf https://doi.org/10.22008/gpub/32631 https://doi.org/https://doi.org/10.1016/j.apenergy.2021.117603 https://doi.org/https://doi.org/10.1016/j.apenergy.2021.117603 scaling the danish national water resources model for a pan-european quasi-3d groundwater resources model introduction geological setting the dk-model methods results discussion conclusions acknowledgements references figures fig. 1 examples of mapped features from the pan-european groundwater resources map (egdi 2022). (a) aquifer types. (b) total depth of aquifers below the surface. the blank areas (grey shading) reflect the countries or regions that did not participate or finalise the grid. denmark is situated within the black square. fig. 2 principal sketches for the layers in the dk-model for jylland, modified from stisen et al. 2019. qs: quaternary sand. qc: quaternary clay. ps: pre-quaternary sand. pc: pre-quaternary clay. fig. 3 example results from a representative cell (cell no. 433_358, for location see fig. 4). (a) representative 2d profile through the dk-model in the cell. (b) extent and calculated mean thicknesses of the saturated and unsaturated part (black bars) of each layer in the cell – here seen in 2d but representing the entire 3d cell (10 × 10 km). fig. 4 calculated mean thicknesses (a) and mean extents (b) for layer 6 (corresponding to valley sand) in 10 × 10 km cells covering the island of fyn. overlaying this are the exact position and thickness of the valley sand layer extracted from the dk-model, shown as a coloured contour map (colour scale in a). the position of the profile in fig. 3a and the cell in fig. 3b are marked on the map. sheldon et al. 2025: geus bulletin 59. 8381. https://doi.org/10.34194/5k9dv133 1 of 30 research article a multidisciplinary biostratigraphic framework for the lower to middle miocene of the norwegian north sea – the siliceous succession of the valhall–hod area emma sheldon1* , karen dybkjær1 , erik skovbjerg rasmussen2 , mimmi oksman3 1department of geo-energy and storage, geological survey of denmark and greenland (geus), copenhagen, denmark; 2department of geophysics and sedimentary basins, geological survey of denmark and greenland (geus), copenhagen, denmark; 3department of glaciology and climate, geological survey of denmark and greenland (geus), copenhagen, denmark abstract a new multidisciplinary biostratigraphic framework, combining dinoflagellate cysts, microfossils, calcareous nannofossils, diatoms and silicoflagellates, is established for the early to middle miocene deep marine clay and siliceous ooze in the southern norwegian sector of the north sea, based on core samples from the valhall and hod hydrocarbon fields. the framework was successfully tested on the equivalent chronostratigraphic level of several wells based on ditch cutting samples. new biostratigraphic events for the danish and norwegian north sea resulting from this study are successfully used to correlate between the valhall and hod areas and supplement published zonation schemes. to our knowledge, this is the first time that diatoms and silicoflagellates from the fine fraction of microfossil samples have been used as correlation tools in the north sea basin. dating of the siliceous/diatomite-rich interval results in a high-resolution (5–15 m intervals) biostratigraphic subdivision. the successful application of the new framework across the valhall and hod areas implies that it could also be useful in a more regional context. the new biostratigraphy enables the correlation of the lithostratigraphic units recently defined for the danish offshore neogene succession to the study area and the correlation of the sequence stratigraphic surfaces defined for the danish sector to the southern norwegian sector. *correspondence: es@geus.dk received: 21 aug 2024 revised: 07 may 2025 accepted: 11 jun 2025 published: 19 dec 2025 keywords: biostratigraphy, diatomite, miocene, north sea, norway abbreviations: fo: first occurence fsst: falling stage systems tract geus: geological survey of denmark and greenland hst: highstand systems tract lo: last occurence lst: lowstand systems tract tst: transgressive systems tract mmct: middle miocene climatic transition mco: miocene climatic optimum prz: partial range zone geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: mette olivarius (geus, denmark) reviewed by: haydon bailey (independent researcher, uk), erik anthonissen (equinor asa, norway) funding: see page 26 competing interests: see page 26 additional files: see page 26 1. introduction a new, multidisciplinary biostratigraphic study of the lower and middle miocene sections of six wells from the valhall and hod fields is presented. these upper cretaceous – danian chalk hydrocarbon fields are located on salt structures in the southernmost part of the norwegian north sea (fig. 1). the miocene succession above the chalk comprises deep-marine clay with a variable content of siliceous ooze. the silica content reaches 50% in some intervals, which are often referred to as diatomite. the focus is on the siliceous ooze or diatomite in some areas, such as the valhall–hod area, because of its hydrocarbon reservoir potential and because its geomechanical properties are critical in connection with well abandonment. two of the six studied well sections, 2/11–12s (hod field) and 2/8–g10a (valhall field), were cored through the lower and middle miocene successions, providing an exceptional and continuous record, unique in the north sea area. the other four studied wells, 2/8–n4, 2/8–v6, 2/8–8 (all from valhall field) and 2/11–1 (in the saddle between the valhall and hod fields), were not cored. the locations of the wells are shown in fig. 2. the purpose of this study is to establish a high-resolution multidisciplinary biostratigraphy for the siliceous or diatomite-rich succession represented by the unique hod and valhall cores. the resulting biostratigraphic framework is then applied to the four non-cored wells using fossil assemblages from ditch cutting samples. https://doi.org/10.34194/5k9dv133 https://orcid.org/0000-0003-4353-8241 https://orcid.org/0000-0002-8420-3379 https://orcid.org/0000-0001-8603-8429 https://orcid.org/0000-0002-8386-516x mailto:es@geus.dk https://creativecommons.org/licenses/by/4.0/deed.en sheldon et al. 2025: geus bulletin 59. 8381. https://doi.org/10.34194/5k9dv133 2 of 30 geusbulletin.org the study was performed combining five biostratigraphic groups: dinoflagellate cysts (dinocysts), microfossils (primarily foraminifera but also including large diatoms and bolboforma), small fraction diatoms, silicoflagellates and calcareous nannofossils, mostly on the same series of closely spaced sediment samples. this is the first time, to our knowledge, that a detailed north sea miocene biostratigraphic study includes small siliceous diatoms and silicoflagellates. 2. geological setting and palaeoclimate the oligocene–miocene transition was characterised by inversion tectonism resulting in shallower waters in the north-eastern part of the north sea basin (ziegler 1990; rasmussen 2009, 2013; knox et  al. 2010). the initial uplift of the southern scandes (fig. 1) re-exposed the present-day norway and central sweden, which formed a low relief landscape at the end of the oligocene (thyberg et al. 2000; løseth & henriksen 2005; gabrielsen et al. 2009). the uplift of the hinterland in the early miocene and the shallowing of the north-eastern north sea basin resulted in progradation of large delta systems from scandinavia (rasmussen et al. 2010; fig. 1). in the northern north sea basin, delta progradation occurred from the west, the shetland platform, coincident with the eastern system (eidvin et al. 2014a). the southern north sea basin was dominated by a coastal plain, and swamp environments formed the margin of a low-relief central european landscape, which was separated from the alps by a foreland basin (fig. 1). during the middle miocene, a major transgression occurred, and the deltas established during the early miocene were flooded. the flooding commenced coincident with a global climatic deterioration (zachos et al. 2001) and the initiation of a new tectonic regime in the north atlantic. huge inversion structures were formed off west norway, and the main phase of uplift of the sole pit structure in the western part of the north sea basin took place (knox et  al. 2010; løseth et  al. 2017 and references therein). iceland also formed at this time (rasmussen et  al. 2008), so branches of the icelandic plume (schoonman et  al. 2017) may also have reshaped the landscape around the northern north sea basin. the late early miocene to middle miocene was also an important phase in the uplift of the carpathian fig. 2 depth map to the top miocene of the valhall and hod structures and locations of the six studied wells (yellow dots). wells circled in red are cored. credit: aker bp. valhall field 2/8-n4 2/8-8 2/8-g10a –1325 –1350 –1375 –1400 –1425 –1450 –1475 –1500 –1525 –1550 –1575 –1600 m 2/8-v6 hod field 2/11-1 2/11-12s 5000 m fig. 1 palaeogeography of the north sea area in the early miocene. the red dot indicates the location of the study area. arrows indicate sediment influx. the valhall/hod area was located in the central part of the basin. north sea sectors are as follows: d: germany. dk: denmark. n: norway. nl: netherlands. uk: united kingdom. modified from rasmussen et al. (2008). 0°e 10°e 50°n 60°n 100 km shetland platform n uk dk d nl valhallhod area southern scandes https://doi.org/10.34194/5k9dv133 http://www.geusbulletin.org/ sheldon et al. 2025: geus bulletin 59. 8381. https://doi.org/10.34194/5k9dv133 3 of 30 geusbulletin.org mountains in central europe (oszczypko 2006). the establishment of the new tectonic regime resulted in accelerated subsidence of the north sea basin, and the deposition of marine mud dominated the middle miocene (koch 1989; rasmussen 2004a, 2004b; rasmussen & dybkjær 2014). our study area was situated in a fully marine, outer shelf to upper bathyal, basin floor setting in a semi-closed basin with long distances to coastlines. deposition was characterised by hemipelagic sedimentation in water depths of between 500 and 1000 m. during the late miocene, continued growth of the carpathian mountains, uplift of the alpine foreland basin and formation of the jura mountains resulted in the formation of a massive new source area in central europe (kuhlemann 2007; fig. 1). in the late miocene, huge, braided river systems supplied the south-eastern north sea basin for the first time (knox et al. 2010). the new central european river system evolved into the so-called eridanos delta system (biljsma 1981; overeem et al. 2001; rasmussen & dybkjær 2014), which began to fill the eastern north sea basin. delta systems sourced from scandinavia also began to prograde into the north-eastern part of the basin during the latest late miocene and reached the central graben area during the messinian stage when they coalesced with the eridanos delta. these delta systems correlate with the nordland group of the norwegian part of the north sea (fig. 3). the climate in the study area was warm-temperate to sub-tropical and humid during the early and middle miocene (utescher et  al. 2009; larsson et  al. 2011; sliwinska et al. 2024). studies on the sdr. vium borehole, jylland, denmark, by larsson et al. (2011), herbert et al. (2020) and sliwinska et al. (2024) indicate mean annual temperatures of around 17–18.5°c on land, mean annual precipitation of c. 750–1750 mm/yr and sea surface temperatures of 23–28°c, although with some fluctuations during the miocene climatic optimum (mco) (c.17–13 ma; e.g. larsson et al. 2011; herbert et al. 2020; sliwinska et al. 2024). at the end of the middle miocene, the global climate deteriorated, a period known as the middle miocene climatic transition (mmct). in the danish and german areas, a decrease in annual temperatures during the serravalian stage has been recognised (utescher et al. 2009; herbert et al. 2020, sliwinska et al. 2024). marked climatic deterioration in the messinian stage at the close of the miocene epoch resulted in the expansion of ice caps on antarctica and probably also  in parts of the northern hemisphere (utescher et al. 2009). 3. lithostratigraphy the neogene deposits in the north sea area include marginal, fluvio-deltaic deposits, shoreface and offshore shelf deposits and basinal deep water hemipelagic and gravity-flow deposits. silicaor diatomite-rich deposits are found locally in basinal areas (including the study area) in the lower and middle miocene parts of the succession. a new lithostratigraphic subdivision of the neogene succession in the danish north sea sector is presented in rasmussen et al. (in press; modified version presented here, fig. 3). in this study, we correlate the new offshore danish lithostratigraphy to the norwegian sector of the north sea. this lithostratigraphy includes the new lower miocene dany formation, which comprises muddy and silty deposits, and the new middle miocene nora formation, which is defined based on its high content of silica or diatomite. the well sections presented in this study include the dany, nora and hodde formations following the new subdivision. according to the norwegian lithostratigraphy for the late paleocene to neogene, the studied lower and middle miocene succession is subdivided into the hordaland group (lark formation) and the nordland group (eidvin et al. 2022; fig. 3). in this study, lithostratigraphic information including diatomite content is available for the two cored wells 2/8–g10a and 2/11–12s, and thus, it has been possible to subdivide the miocene succession into the lithostratigraphic units defined in the danish sector. an attempt at a similar subdivision in the non-cored wells is made, based on gamma log responses. 4. sequence stratigraphic framework in the danish and southern norwegian sectors, eight depositional sequence boundaries are found within the miocene succession (rasmussen et al. 1996, rasmussen 2004b, 2017). the boundaries are defined on a combination of studies of outcrop sections onshore denmark, borehole logs and cores and seismic data, including high-resolution shallow seismic, multichannel, and for the central graben area, 3d seismic data (rasmussen 1996; rasmussen 2004b, 2017, dybkjær et al. 2021). the eight sequence boundaries confine seven fully developed sequences, named b, c, d1, d2, e, f1 and f2 (rasmussen 2004b, 2017). the two lowermost sequences, b and c, and the f2 sequence include all four systems tracts (lst: lowstand systems tract, tst: transgressive systems tract, hst: highstand systems tract and fsst: falling stage systems tract), whereas the upper lower miocene to lower upper miocene sequences only have a two-fold subdivision (tst and hst). the latter was due to increased subsidence of the north sea basin during the middle part of the miocene outpacing eustatic sea-level fall (rasmussen 2004b, 2017). the seismic surfaces can be seen for the 2/8–g10a and 2/11–125 wells in figs 5 and 6 and for all wells in the supplementary files s1–s12. https://doi.org/10.34194/5k9dv133 http://www.geusbulletin.org/ sheldon et al. 2025: geus bulletin 59. 8381. https://doi.org/10.34194/5k9dv133 4 of 30 geusbulletin.org fi g. 3  t he n ew li th os tr at ig ra ph ic s ub di vi si on fo r th e n eo ge ne s uc ce ss io n in th e d an is h se ct or o f t he n or th s ea (m od ifi ed fr om r as m us se n et  a l., in p re ss ) w ith a d as he d re d bo x sh ow in g th e st ud ie d in te rv al . n ot e th e pr es en ce o f d ia to m ite in th e m id dl e m io ce ne , r ef er re d to a s th e n or a fo rm at io n (s ol id r ed b ox ), an d so m e m in or o cc ur re nc es in th e lo w er m io ce ne s uc ce ss io n w ith in th e d an y fo rm at io n (r ed a rr ow s) . t he in te rv al s co ve re d by th e tw o co re d w el l s ec tio ns (2 /1 1– 12 s an d 2/ 8– g 10 a) a re s ho w n. t he ti m es ca le o f r affi e t a l. (2 02 0) is u se d in th is fi gu re . a qu it an ia n bu rd ig al ia n la ng hi an se rr av al lia n to rt on ia n m es si ni an za nc le an pi ac en zi an g el as ia n c al ab ri an neogenequaternary miocene middleupper lower pliocene pleistocene h ol oc en e 0 5 10 15 20 måde group ribe group lark formation nordland group hordaland group g ra m f m ø rn hø j f m h od de f m a rn um f m ba st ru p fm v ej le f jo rd f m bi llu nd f m kl in ti ng ho ve d fm fl oo dp la in s ed im en t d el ta s an d o �s ho re d ia to m it e sw am p fl uv ia l s an d/ gr av el f2 f1 e d 2 d 1 c b a qu i10 0 a qu i2 0 0 bu rd 10 0 bu rd 20 0 la ng 10 0 to rt 10 0 to rt 20 0 m es s1 0 0 sl op e sa nd o �s ho re g ra vi ty -� ow s an d o �s ho re m ud o �s ho re s an d d ee p m ar in e m ud sh or ef ac e sa nd g ra m f m em m a fm el in f m emma fm ø rn hø j f m h od de f m g ra m f m ø rn hø j f m g 10 a 12 s h od de f m a rn um f m o dd er up f m n or a fm o dd er up f m kl in ti ng ho ve d fm d an y fm d an y fm m ar bæ k fm li lle jo hn m b m ar bæ k fm v ag n fm h . o bs cu ra (h .o .) s. ar m ag ed do ne ns is (s .a .) m . c ho an op ho ru m (m .c .) b. p lio ce ni cu m (b .p .) i. m ul tip le xu m (i. m .) g . v er ri cu la (g .v .) a . a nd al ou sie ns e (a .a .) a . u m br ac ul a (a .u .) u . a qu ae du ct um (u .a .) l. tr un ca tu m (l .t. ) c . c an th ar el lu s ( c .c .) e. in sig ne (e .i. ) c . a ub ry ae (c .a u. ) c . g al ea (c .g .) s. ha m ul at um (s .h .) t. pe la gi ca (t .p .) c . a m ic ul um (c .a m .) h om ot ry bl iu m sp p. (h .) o dd er up f m o dd er up f m v ag n fm lu na f m a ge w e (m a) period ep oc h a ge /s ta ge n or w eg ia n li th ost ra ti gr ap hy c en tr al g ra be n d in oc ys t zo na ti on (d yb kj æ r & pi as ec ki 2 0 10 ) ri ng kø bi ng – f yn h ig h n or w eg ia n d an is h ba si n/ on sh or e d en m ar k se qu en ce s (r as m us se n 20 17 ) se qu en ce s (in fo rm al ) https://doi.org/10.34194/5k9dv133 http://www.geusbulletin.org/ sheldon et al. 2025: geus bulletin 59. 8381. https://doi.org/10.34194/5k9dv133 5 of 30 geusbulletin.org 5. absolute dating absolute dating by palaeomagnetic stratigraphy, radiometric dating or sr isotope stratigraphy of the miocene succession in the north sea basin is rare. deeper levels were usually targeted for hydrocarbon exploration, while the younger ‘overburden’ was traditionally considered uninteresting and therefore only rarely cored. therefore, the absolute ages of dinocyst and microfossil events within the miocene succession in the north sea basin as shown by powell (1992), munsterman & brinkhuis (2004), louwye et  al. (2007), dybkjær & piasecki (2010), köthe (2012), king (1989, 2016), munsterman et al. (2019) and dybkjær et al. (2019) are usually based on data from areas outside the north sea basin, where these events are found in wells where absolute dating has been carried out (e.g. haq et al. 1987; de verteuil & norris 1996; de verteuil 1997; williams et al. 2004 and references therein). also, correlation with other microfossil groups (e.g. nannofossils) and the global sea-level changes have been used to date the dinocyst and microfossil events recorded in the north sea basin (e.g. dybkjær & piasecki 2010; king 2016; munsterman et al. 2019). however, the absolute dating of specific events varies from one reference to another, due to the diachronicity of first and last appearance datums and uncertainties of the datings. the most comprehensive sr-isotope study of the miocene north sea basin is that of eidvin et al. (2014b) from onshore denmark. the results of that study generally supported the ages of the dinocyst zones of dybkjær & piasecki (2010) in the early miocene but also documented the uncertainty of using sr-isotopes for dating this stratigraphic level, especially the late middle to late miocene part of the succession. comparing the dinocyst event and zonation scheme of dybkjær & piasecki (2010, their fig. 6) with that of king (2016, their fig. 18) clearly reflects this uncertainty in chronostratigraphic correlation. similarly, correlation between dinocyst and microfossil events and zones also differs in the two publications (e.g. dybkjær & piasecki 2010, their fig. 6; king 2016, their fig. 21). correlation between miocene nannofossil and microfossil zones is also problematic, as seen in and explained by king (2016), compare their figs. 21 and 27. absolute dating has not been carried out on the studied wells. due to this absence of an absolute age model, the events and biozonations for the five studied microfossil groups are presented against established zonations (fig. 4), sample depth, sequence boundaries, lithostratigraphy and the gamma log, instead of against a timescale (ma), see figs 5, 6 and supplementary files s1–s12. 6. previous studies the majority of biostratigraphic work that has been carried out on the neogene succession of the north sea area over the past 50 years is a direct result of extensive hydrocarbon exploration at deeper levels. palynological studies on the miocene succession in the north sea include piasecki (1980), strauss & lund (1992), powell (1992), head (1996), louwye et  al. (1999), louwye (2002), dybkjær & rasmussen (2000, 2007), strauss et  al. (2001), de schepper et  al. (2004, 2009), dybkjær (2004a, 2004b), munsterman & brinkhuis (2004), schiøler (2005), köthe & piesker (2007), louwye et al. (2007), louwye & laga (2008), louwye & de schepper (2010), dybkjær & piasecki (2010), dybkjær et al. (2012), köthe (2012), eidvin et al. (2014a, 2014b), śliwinska et al. (2014), king (2016), de schepper & mangerud (2017), grøsfjeld et al. (2019), dybkjær et al. (2021) & sliwinska et al. (2024). microfossil studies of the miocene of the north sea area are also numerous (von daniels & spiegler 1977; doppert et al. 1979; doppert 1980; spiegler & von daniels 1991; spiegler 1999; king 1983, 1989, 2016; gradstein et al. 1988; gradstein & backstrom 1996; laursen & kristoffersen 1999; eidvin et al. 1999; kaminski & gradstein 2005; rundberg & eidvin 2005; eidvin & rundberg 2007; anthonissen 2012; fox et al. 2018). published neogene calcareous nannofossil studies for the north sea area are not common due to the siliciclastic nature of much of the neogene section and the successful application of high-resolution studies of other biostratigraphic disciplines. neogene nannofossil studies of the broader north atlantic region include müller (1976), steinmetz (1979), gartner (1992), de kaenel et  al. (2017), boesiger et  al. (2017) and bergen et  al. (2017). the global nannofossil ‘nn’ zonation of martini (1971) is still widely used and is correlated with other nannofossil zonations in young et al. (1994) and young (1998). for the north sea area, the use of diatoms for palaeogene and neogene biostratigraphy was studied by mitlehner (2019). several biostratigraphic studies of upper oligocene and lower miocene north sea successions using large, pyritized diatoms in conjunction with other microfossils have also been published. king (1983, 2016) focused on the cainozoic micropalaeontological biostratigraphy of the north sea and adjacent areas. laursen and kristoffersen (1999) studied the miocene successions of onshore denmark using foraminifera, and dybkjær et al. (2012) concentrated on the oligocene–miocene boundary in the eastern north sea basin using dinocyst stratigraphy, micropalaeontology and δ13c-isotope data. eidvin and rundberg (2001) focused on the late cainozoic stratigraphy of the northern north sea, and https://doi.org/10.34194/5k9dv133 http://www.geusbulletin.org/ sheldon et al. 2025: geus bulletin 59. 8381. https://doi.org/10.34194/5k9dv133 6 of 30 geusbulletin.org anthonissen (2012) compiled an integrated miocene biostratigraphy for the northeastern north atlantic. small siliceous diatoms are not generally used in the north sea area for routine biostratigraphy as their minute size results in them being washed through standard sieves, and their delicate structures are easily destroyed by the harsh preparation techniques used in industry. a diatom zonation scheme for the north sea does not exist. however, studies with potential use in biostratigraphy in the neogene of the north sea area include eidvin et al. (1998), thyberg et al. (1999) and sheldon et al. (2018). siliceous diatom studies from the wider high northern latitudes that are useful for correlation with this study include those of schrader & fenner (1976), koç & scherer (1996) and suto (2006) from the norwegian and iceland seas, dzinoridze et al. (1979) from the norwegian basin and baldauf (1985) from the rockall plateau. silicoflagellates only comprise a small percentage of the siliceous component of marine sediments and therefore have limited biostratigraphic use. a silicoflagellate zonation scheme for the north sea does not exist. martini & muller (1976), locker & martini (1989), ciesielski et al. (1989) and amigo (1999) investigated miocene successions in the norwegian-greenland sea and the iceland-rockall plateau areas to the north of the study area. 7. zonation schemes in this study the biostratigraphic zonations used in this study are the north sea dinocyst zonation of dybkjær & piasecki (2010), the north sea microfossil zonations of king (1989, 2016), the global calcareous nannoplankton zonation of martini (1971), the norwegian sea diatom zonation of schrader & fenner (1976) and the norwegian sea silicoflagellate zonation of locker & martini (1989), fig. 4. in their study of the neogene of the iceland sea, koç & scherer (1996) revised the biostratigraphy of schrader & fenner (1976). in this study, however, we revert to the zonation of schrader & fenner (1976) due to the similarity of diatom assemblages therein with those from the valhall–hod area. the zonation schemes follow a chronostratigraphy that was current when the zonation was published. the north sea dinocyst zonation of dybkjær & piasecki (2010) follows the chronostratigraphy of lourens et al. (2004). the north sea microfossil zonations of king (1989, 2016) follow the chronostratigraphy of hilgen et al. (2012) in king (2016). the global calcareous nannoplankton zonation of martini (1971) is correlated with the chronostratigraphy of raffi et  al. (2020). the norwegian sea diatom zonation of schrader & fenner (1976) follows the chronostratigraphy of berggren (1972), and the norwegian sea silicoflagellate zonation of locker & martini (1989) follows the chronostratigraphy of berggren et al. (1985). it is beyond the scope of this paper to attempt an up-to-date chronostratigraphic correlation of the five zonation schemes. dinocyst taxonomy follows the ‘lentin & williams index’ (williams et al. 2017). microfossil taxonomy follows that used in king (1989, 2016) and young et  al. (2024b). nannofossil taxonomy is based on young et al. (2024a). diatom taxonomy follows that of schrader & fenner (1976) and barron (1985), and silicoflagellate taxonomy follows that of perch-nielsen (1985) and locker & martini (1989). 8. materials and methods cores from the 2/11–12s (hod field) and 2/8–g10a (valhall field) wells were sampled with a spacing of approximately 4 to 5 m. ditch cutting samples were taken from the core gap in 2/8–g10a. for the non-cored wells 2/8– n4, 2/8–v6, 2/8–8 (all valhall field) and 2/11–1 (in the saddle between the valhall and hod fields), ditch cutting samples were taken every 10 m (table 1). the position of the analysed samples in each well is shown in figs 5, 6 and supplementary files s1–s12. most of the samples were analysed for palynology (dinocysts), microfossils (large fraction) and siliceous microfossils (small fraction: diatoms and silicoflagellates). only around ten samples per well were selected from minor calcareous-rich intervals and analysed for nannofossils as they are not routinely used for biostratigraphy in the miocene of the north sea. 8.1. preparation methods the sample preparation methods for each biostratigraphic discipline are described below. all sediment samples were processed in the stratigraphic laboratory at the geological survey of denmark and greenland (geus). 8.1.1. palynology approximately 20 g of sample was dried and crushed until all particles were <2 mm in size. dissolution of table 1 number of samples for each discipline for each well. well 2/11-12s well 2/8-g10a well 2/8-v6 well 2/8-n4 well 2/8-8 well 2/11-1 palynology (dinocysts): 40 40 co, 15 dcs 27 25 49 32 microfossils: 47 39 co, 13 dcs 27 25 51 34 diatoms & silicoflagellates: 36 39 co, 13 dcs 27 25 51 34 nannofossils: 10 10 co 10 11 11 11 co: core samples. dcs: ditch cutting samples. the samples from 2/11-12s are all from core material. those from 2/8-g10a are mostly from core material and the remaining well samples are ditch cutting samples. https://doi.org/10.34194/5k9dv133 http://www.geusbulletin.org/ sheldon et al. 2025: geus bulletin 59. 8381. https://doi.org/10.34194/5k9dv133 7 of 30 geusbulletin.org fi g. 4  b io st ra tig ra ph ic z on at io ns u se d in th is s tu dy . d in oc ys t z on at io n of d yb kj æ r & p ia se ck i ( 20 10 ), na nn of os si l z on at io n of m ar tin i ( 19 71 ), m ic ro fo ss il zo na tio ns o f k in g (1 98 9, 2 01 6) , d ia to m z on at io n of s ch ra de r & fe nn er (1 97 6) & th e si lic ofl ag el la te z on at io n of l oc ke r & m ar tin i ( 19 89 ). re fe r to in di vi du al r ef er en ce s fo r tim es ca le s ap pl ie d. ep oc h st ag e d in o� ag el la te s zo ne a ge (m a) a ge (m a) a ge (m a) a ge (m a) a ge (m a) 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 11 12 13 14 15 16 17 18 19 20 21 22 23 10 11 12 13 14 15 16 17 18 19 20 21 22 23 13 14 15 16 17 18 19 20 21 22 9 10 11 12 13 14 15 16 17 18 19 20 21 ep oc h st ag e n an no fo ss ils zo ne ep oc h st ag e m ic ro fo ss ils n sb z on e n s zo ne ep oc h d ia to m s pa rt ia l r an ge z on e ep oc hsi lic o� ag el la te s zo ne miocene serravalian langhian burdigalian aquitaniantortonian miocene miocene serravalian langhian burdigalian aquitaniantortonian n s4 0 n s3 9 n s3 8 n s3 7 n s3 6 n s3 5 bb cc aaa n s3 4 th al as sio sir a fr ag a n itz ec hi a m al ei nt er pr et ar ia c os oi no di so us v ig ila ns rh iz os en ia n or w eg ic a sy ne dr a jo us ea na ps eu do di m er og ra m m a el eg an s c os ci no di sc us p lic at us d en tic ul a hy al in a u pp er c . t ria ca nt ha h ia tu s p. c irc ul us a pi cu la ta lo w er c . t ria ca nt ha n . n av ic ul a rh iz os ol en ia b ul bo sa 9101112 b 12 c 12 a n s3 3 n n 8 n n 7 n n 6 n n 5 n n 4 n n 3 n n 2 n n 1 serravalian langhian burdigalian aquitaniantortonian early miocenemiddle miocene early miocenemiddle miocenelate miocene a . u m br ac ul a g . v er ric ul a a . a nd al ou sie ns e h . o bs cu ra u . a qu ae du ct um l. tr un ca tu m c . a ub ry ae e. in sig ne c . c an th ar el lu s s. h am ul at um t. p el ag ic a c . a m ic ul um h om ot ry bl iu m s pp . c . g al ea https://doi.org/10.34194/5k9dv133 http://www.geusbulletin.org/ sheldon et al. 2025: geus bulletin 59. 8381. https://doi.org/10.34194/5k9dv133 8 of 30 geusbulletin.org fi g. 5  c ap tio n on n ex t p ag e. •• a bu nd an t • c om m on ° c on si st en t w el l n am e: 2 /8 g 10 a fi g 5a 14 60 14 50 14 70 14 80 14 90 15 0 0 15 10 15 20 15 30 15 40 15 50 15 60 15 70 15 80 15 90 gamma log cores li th ost ra ti gr ap hy chronostratigraphy d in o� ag el la te c ys t e ve nt s samples north sea cenozoic microfossils m ic ro fo ss il ev en ts sb e m fs d 2 m rs e nordland gp. måde group lower langhianupper langhian middle miocene dany fm. nora fm. ribe group ørnhøj fm. hodde fm. o�shore denmark o�shore norway hordaland group lark fm. * ** ** * ? m .m . se r. a . a nd . ? ? ? ? ? labyrinthodinium truncatumunipontidinium aquaeductum nannofossil zone nannofossil samples n an no fo ss il ev en ts diatom p.r.z. d ia to m e ve nt s silico�agellate zone si lic o� ag el la te e ve nt s s. a �. k itt on ia nu s d . h us te dt ii, o . g em m at a, sy ne dr a sp p. , s . g ru no w ii, d en tic ul op sis s pp . • h ig h ab un d. & d iv er si ty d ia to m a ss em bl ag es a . i ng en s, d . s m ith ii p. e lo ng at a, d . h ya lin a c . b ih ar en sis t. n itz sc hi od es •• r. w ic om ic oe ns is d en tic ul op sis s pp . • , d . h us te dt ii • a . i ng en s ° s. h or rid us •• s. g ru no w ii °, sy ne dr a sp p. (i n� ux ) d . c ru x t. fragarhizosolenia bulbosadenticulopsis hyalina lower corbisema triacanthaupper corbisema triacantha r. m io ce ni ca b. d io do n b. d io do n, o . s pe cu lu m c . t ria ca nt ha c . b ih ar en sis c . p ep lu m north sea cenozoic microfossils c . a ng ul io � ci na lis t. q ua dr ilo ba tu s b. p la ty re tic ul at a a . g . s ta es ch ei , l. si nu os um , e. in �a tu m , c . s ub co ni cu s, l. p er eg rin a e. in �a tu m t. tr ilo bu s, c . d . p ee le ns is, g . e ol ab ia cr as sa ta , p. n an a, c . c ip er oe ns is undi� undi�erentiated nsb12a nsb11 undi�erentiated ns36c ns36a-b h . s ci ss ur a (o cc ), d . e xi lis (o cc ), h . w al tr an s ( oc c) , d . d isc iss us (o cc ) d . c au li� or is, h . b ip un ct a, c . � or id an us , c . p el ag ic us , h . c ar te ri d . d ru gi i ( oc c) , d . d e� an dr ei (o cc ) h . b ip un ct a s. n eo ab ie s u . a qu a. , i m p. s p. a w & k 19 96 •, p. m io ca en ic um a . a nd al ou sie ns is, o . t eg ill at um im pa gi di ni um s p. a w re nn & ko ki no s 19 96 •, i. la cr ym os a, o . e iri ki an um h . t ec ta ta • h . t ec ta ta u . a qu ae du ct um c . a ub ry ae i. ta bu la ta d . h ya lin a t. fr ag a n n 4 n n 45 sequence stratigraphy dino�agellate cyst zone ki ng 19 89 ki ng 20 16 u nd i� c . p as ., c . p ou l., c . p la c. (o cc ) a . a nd al ou sie ns is https://doi.org/10.34194/5k9dv133 http://www.geusbulletin.org/ sheldon et al. 2025: geus bulletin 59. 8381. https://doi.org/10.34194/5k9dv133 9 of 30 geusbulletin.org fi g. 5  ( co nt in ue d) s um m ar y di ag ra m fo r th e 2/ 8– g 10 a w el l w ith b io st ra tig ra ph ic z on es , p re vi ou sl y pu bl is he d ev en ts (b la ck t ex t) a nd n ew e ve nt s (r ed t ex t) , f or fu ll fo ss il na m es s ee s up pl em en ta ry f ile s 1. l ith os tr atig ra ph y fo r th e n or w eg ia n se ct or fr om e id vi n et  a l. (2 02 2) , o ns ho re d en m ar k fr om r as m us se n et  a l. (2 01 0) a nd o ff sh or e d en m ar k fr om r as m us se n et  a l. (in p re ss ). se qu en ce s tr at ig ra ph y fr om r as m us se n (2 00 4a ; 20 17 ) a nd d yb kj æ r et  a l. (2 02 1) . g am m a lo g c/ o ak er b p. c or ed in te rv al s ar e in di ca te d. t he m ai n sa m pl e co lu m n in di ca te s sa m pl es fo r di no cy st s, m ic ro fo ss ils , d ia to m s an d si lic ofl ag el la te s. n an no fo ss il sa m pl es a re in di ca te d se pa ra te ly . i n th e ‘c or es ’ c ol um n, c or e sa m pl es a re m ar ke d w ith a d as hdo t sy m bo l, an d di tc h cu tt in g sa m pl es a re m ar ke d w ith a d as h. c hr on os tr at ig ra ph y of r affi e t a l. (2 02 0) , d in oc ys t z on at io n of d yb kj æ r an d pi as ec ki (2 01 0) , m ic ro fo ss il zo na tio ns o f k in g (1 98 9; 2 01 6) , d ia to m z on at io n of s ch ra de r an d fe nn er (1 97 6) , s ili co fla ge lla te z on at io n of l oc ke r an d m ar tin i ( 19 89 ) a nd n an no fo ss il zo na tio n of m ar tin i ( 19 71 ). *: l at e m io ce ne , * *: la te t or to ni an , * ** : h ys tr ic ho sp ha er op si s ob sc ur a zo ne . f m : f or m at io n. m .m . a nd m . m ic oe ne : m id dl e m io ce ne . s er .: se rr av al lia n. l an g. : l an gh ia n. p .r .z .: pa rt ia l r an ge z on e. u .: u pp er . l w r. : l ow er . a qu i.: aq ui ta ni an . b ur d. : b ur di ga lia n. u nd iff .: u nd iff er en tia te d. lower lang. m. miocene l. truncatum l. tr un ca tu m ° 17 0 0 17 10 17 30 17 40 17 20 sb d 1 lwr. aqui.u. aqui. – lwr. burd lw r. s. h am . – c . a m ic . st ic to di sc us a �. k itt on ia nu s, o . g em m at a h ig h ab un da nc e & di ve rs ity d ia to m as se m bl ag es undi� p. se m in ud a, a . a llo rg ei , a . a em ul an s s. h am ul at um t. p el ag ic a (o cc ) e. b ur di ga le ns is e. b ur di ga le ns is c hi ro pt er id iu m s pp . ?c . g al ea n sb 9 n s3 4 ?h om ot r. sp p. 16 80 16 90 upper s. hamulatum p. e lo ng at a undi� nn3-4 ns35a a . g . s ta es ch ei e. in sig ne t. ro ta d . c la do id es w el l n am e: 2 /8 g 10 a fi g 5b 15 90 16 0 0 16 10 16 20 16 30 16 40 16 50 16 60 16 70 gamma log cores li th ost ra ti gr ap hy chronostratigraphy d in o� ag el la te c ys t e ve nt s samples north sea cenozoic microfossils m ic ro fo ss il ev en ts sb d 2 m fs d 1 m rs d 1 lower miocene burdigalian dany fm. ribe group o�shore denmark o�shore norway hordaland group lark fm. c. canth.cousteaudinium aubryae nannofossil zone n an no fo ss il ev en ts diatom p.r.z. d ia to m e ve nt s silico�agellate zone si lic o� ag el la te e ve nt s r. m ar ga rit al im ba ta , r. ro bu st at a, r. m ar ga rit al im ba ta t. n itz sc hi od es ••• r. g al id a, s . o ss ifo rm is r. norwegica c. vigilans n. maleinterpretariathalassiosira fraga lower corbisema triacantha o . s pe cu lu m , d . c ru x c . t ria ca nt ha s. h or rid us ••• north sea cenozoic microfossils u . t en ui pu st ul a, g . z ea la nd ic a, g . p ra es ci tu la a . w ol te rs to r� g . z ea la nd ic a t. b ra dy i, n . c f. gr an os um c . c on tr ar ia , g . o bl iq uu s u . t en ui pu st ul a, g . e ol ab ia cr as sa ta (o cc ) g . p ra es ci tu la nn4 nsb10 ns35b d . c au li� or is, s. h et er om or ph us (o cc ) h . a m pl ia pe rt a h . w al be rs do rf en sis (o cc ) h . a m pl ia pe rt a p. m io ca en ic um c . p ou lse ni i c . a ub ry ae c . c an th ar el lu s e. in sig ne sequence stratigraphy dino�agellate cyst zone e. in sig ne ki ng 19 89 ki ng 20 16 r. m ar yl an ic us (o cc ) h . m ed ite rr an ea (o cc ) d . e m bl em at ic us , c . p el ag ic us , h . c ar te ri, s . n eo ab ie s c . � or id an us d . e m bl em at ic us nannofossil samples •• a bu nd an t • c om m on ° c on si st en t https://doi.org/10.34194/5k9dv133 http://www.geusbulletin.org/ sheldon et al. 2025: geus bulletin 59. 8381. https://doi.org/10.34194/5k9dv133 10 of 30 geusbulletin.org fi g. 6  s um m ar y di ag ra m fo r t he 2 /1 1– 12 s w el l w ith b io st ra tig ra ph ic z on es , p re vi ou sl y pu bl is he d ev en ts (b la ck te xt ) a nd n ew e ve nt s (r ed te xt ), fo r f ul l f os si l n am es s ee s up pl em en ta ry f ile s 2. l ith os tr at ig ra ph y fo r t he n or w eg ia n se ct or fr om e id vi n et  a l. (2 02 2) , o ns ho re d en m ar k fr om r as m us se n et  a l. (2 01 0) a nd o ff sh or e d en m ar k fr om r as m us se n et  a l. (in p re ss ). se qu en ce s tr at ig ra ph y fr om r as m us se n (2 00 4a ; 2 01 7) a nd d yb kj æ r et  a l. (2 02 1) . g am m a lo g c/ o ak er b p. c or ed in te rv al s ar e in di ca te d. t he m ai n sa m pl e co lu m n in di ca te s sa m pl es fo r di no cy st s, m ic ro fo ss ils , d ia to m s an d si lic ofl ag el la te s. n an no fo ss il sa m pl es a re in di ca te d se pa ra te ly . co re s am pl es a re m ar ke d w ith a d as h w it h a do t, an d di tc h cu tt in g sa m pl es a re m ar ke d w ith a d as h. c hr on os tr at ig ra ph y of r affi e t a l. (2 02 0) , d in oc ys t z on at io n of d yb kj æ r a nd p ia se ck i ( 20 10 ), m ic ro fo ss il zo na tio ns o f ki ng (1 98 9, 2 01 6) , d ia to m z on at io n of s ch ra de r an d fe nn er (1 97 6) , s ili co fla ge lla te z on at io n of l oc ke r an d m ar tin i ( 19 89 ) a nd n an no fo ss il zo na tio n of m ar tin i ( 19 71 ). fm : f or m at io n. u nd iff .: u nd iff er en tia te d. im pa gi di ni um s p. a w re nn & ko ki no s 19 86 • s. h am ul at um , u . a qu ae du ct um gamma log cores li th ost ra ti gr ap hy chronostratigraphy d in o� ag el la te c ys t e ve nt s samples north sea cenozoic microfossils m ic ro fo ss il ev en ts o�shore denmark o�shore norway nannofossil zone n an no fo ss il ev en ts diatom p.r.z. d ia to m e ve nt s silico�agellate zone si lic o� ag el la te e ve nt s north sea cenozoic microfossils sequence stratigraphy dino�agellate cyst zone ki ng 19 89 ki ng 20 16 w el l n am e: 2 /1 1 12 s f ig 6 14 80 14 90 15 0 0 15 10 15 20 15 30 15 40 15 50 15 60 15 70 15 80 15 90 16 0 0 16 10 16 20 16 30 16 40 16 50 16 60 16 70 16 80 sb f m fs e m rs e sb e m fs d 2 sb d 2 m fs d 1 m rs d 1 nordland gp. lower miocene burdigalianlower langhianupper langhianserravallian middle miocene dany fm. nora fm.hodde fm.ørnhøj fm. måde group ribe group hordaland group lark fm. s. h am . g . ve rr ic ul a a . a nd . c . ca nt h.cousteaudinium aubryaelabyrinthodinium truncatumunipontidinium aquaeductum undifferentiated undifferentiated nn3-4nn4undi� undi� ns36c-38ns38-39 ns36a-b ns35ans35b-c nsb12c nsb12a-b nsb10nsb11 sy ne dr a sp p. d . c ru x d . c ru x o . s pe cu lu m h ig h ab un . & d iv . d ia to m a ss em ., d . h us te dt ii, s . g ru no w ii, d en tic ul op sis s pp . • o . g em m at a d . s m ith ii t. n itz sc hi od es •• a . i ng en s d . h ya lin a, p. e lo ng at a d en tic ul op sis s pp . • •, d . h us te dt ii r. m ar yl an di cu s s. g ru no w ii °, a . i ng en s, s. h or rid us •• s. h or rid us •• t. n itz sc hi od es •• o . g em m at a h ig h ab un da nc e an d di ve rs ity d ia to m as se m bl ag es c . � or id an us h . b ip un ct a, h . a m pl ia pe rt a, d . l en tic ul at a, r. p se ud ou m bi lic us h . s ci ss ur a, d . e m bl em at ic us , d . d ru gg ii, d . c au li� or is, d . e xi lis (o cc ), d . m oo re i ( oc c) , h . w al be rs do rf en sis (o cc ) s. p ro ce ru s ( oc c) , s. h et er om or ph us , s . c on ic us h . m ed ite rr an ea (o cc ), d . p et al lif . ( oc c) , s . a po xi s ( oc c) u . k in gi , b . p la ty re tic ul at a u . k in gi c . a ng ul io � ci na lis , t. q ua dr ilo ba tu s s. d isj un ct a, a . g . s ta es ch ei , l . s in uo su m , c . s ub co ni cu s l. p er eg rin a, c . d . p ee le ns is g . e ol ab ia cr as sa ta , t. tr ilo bu s, e. in �a tu m s. d isj un ct a u . t en ui pu st ul at a, n on io n cf . g ra no su m e. in �a tu m c . c on tr ar ia , g . p ra es ci tu la , g . z ea la nd ic a t. b ra dy i, a . w ol te rs do r� g . o bl iq uu s g . z ea la nd ic a u . t en ui pu st ul a g . p ra es ci tu la g . v er ric ul a, c . p ou lse ni i ( oc c) c . p as sio a . a nd al ou sie ns is c . p as sio , i . l ac ry m os a p. m io ca en ic um , c . p la ca nt hu m o . e iri ki an um , h . t ec ta ta • h . t ec ta ta u . a qu ae du ct um c . a ub ry ae i. ta bu la ta l. tr un ca tu m p. m io ca en ic um c . p ou lse ni i e. in sig ne c . a ub ry ae c . c an th ar el lu s e. in sig ne s. h am ul at um , ( oc c) t. ro ta p. n an a, c . c ip er oe ns is h . b ip un ct a d . d ru gg ii, d . c au li� or is, s. h et er om or ph us denticulopsis hyalina d . h ya lin a, r. m ar ga rit al im ba ta (o cc ) r. a ng ul at a (o cc ) r. w ic om ic oe ns is sy ne dr a sp p. (i n� ux ) e. in sig ne t. fragarhizosolenia bulbosa c . p el ag ic us , h . c ar te ri d . e m bl em at ic us , h . s ci ss ur a s. p un ic eu s, d . d e� an dr ei s. p un ic eu s h . w al tr an s c . � or id an us , h . a m pl ia pe rt a, r. p se ud ou m bi lic us , h . c ar te ri, c . p el ag ic us a . g . s ta es ch ei nannofossil samples •• a bu nd an t • c om m on ° c on si st en t https://doi.org/10.34194/5k9dv133 http://www.geusbulletin.org/ sheldon et al. 2025: geus bulletin 59. 8381. https://doi.org/10.34194/5k9dv133 11 of 30 geusbulletin.org carbonates was carried out using 1m hcl until bubbling ceased, followed by 5m hcl for 24 h. this was followed by treatment with a mild solution of citric acid heated to 70°c. the silicate fraction was dissolved using cold hf (40%) for a minimum of 6 days, followed by treatment with a mild solution of citric acid heated to 70°c. the acid treatment was followed by brief oxidation with concentrated hno3 and koh 5% and by heavy liquid separation with znbr (2.3 g/ml). each step was followed by filtration through an 11 µm nylon net. a final filtration through a 20 µm nylon net was carried out before mounting the acid-resistant organic particles in glycerin jelly on glass slides. the dinocyst slides were examined using a leica dm2000 normal light microscope at 400x and 1000x magnification. a minimum of 200 dinocysts were identified to species level. finally, the slide used for counting and an additional slide were scanned to include rare dinocyst taxa. 8.1.2. microfossils (large fraction) approximately 50 g of sample was boiled to disaggregate it, and a little washing-up liquid was added to help remove drilling mud. the sediment was washed through a 63 µm sieve and dried at 60°c. the residues were sieved into >250 µm, >100 µm and >63 µm fractions. the larger fractions were picked for microfossils (primarily foraminifera, but also large diatoms and bolboforma), and the >63 µm fraction was scanned. the microfossils were examined using a leica m205c microscope and semi-quantitative counting. 8.1.3. calcareous nannofossils nannofossil smear slides were prepared using the simple smear slide technique described in bown & young (1998). the prepared slides were examined using a leica dm2500p light microscope under x1000 magnification with cross-polarised light. several slide traverses were analysed, and simple presence–absence recording was undertaken. 8.1.4. siliceous microfossils (small fraction) the siliceous microfossil (small fraction) preparation method used in this study was developed at geus. hydrogen peroxide was used to remove the organic material. the sample was then boiled for several hours and cooled, and then, a small amount of 10% hcl was added. the sample was then cleaned with distilled water and allowed to rest, several times over, for a few days. microspheres were added and the mixture pipetted onto a glass coverslip and dried overnight. the coverslip was mounted onto a glass slide with naphrax and heated on a hot plate to set. the prepared slides were examined using a leica dm2500p light microscope under x1000 magnification. for the two cored wells, the whole of each slide was analysed, and all diatoms and silicoflagellates were counted. for the non-cored wells, the five ‘longest traverses’ of the circular slide were counted. 9. results high-quality biostratigraphic data were produced from the two cored wells, 2/11–12s and 2/8–g10a; in that cored material is not subjected to harsh mechanical and chemical processes that can destroy fossil assemblages. the data from the ditch cutting samples from the 2/8–n4, 2/8– v6, 2/8–8 and 2/11–1 wells and from the core gap in 2/8– g10a are also generally of good quality, although caved dinocysts and microfossils occur. natural processes such as reworking and fluctuations in abundance and diversity of the dinocysts and microfossils due to climatic or palaeoenvironmental changes also influence the data. the multidisciplinary biostratigraphic study of the lower and middle miocene succession of the 2/8–g10a and 2/11–12s cored sections resulted in the successful application of established regional dinocyst and microfossil biozonations (dybkjær & piasecki 2010; king 1989, 2016) and the global nannofossil zonation of martini (1971), in addition to the recognition of several new key bioevents that appear to have stratigraphic potential (figs. 4–8, supplementary files s1, s2, s13). the biozones and new events identified in the cored wells were tested on the non-cored sections with a high degree of success and are discussed here (see also fig. 8 and supplementary files s3–s13). to our knowledge, diatoms (small fraction) and silicoflagellates have not previously been applied to routine biostratigraphy in the north sea and are therefore discussed in some detail here. fo denotes the first (oldest) stratigraphic occurrence in this study, and lo denotes the last (youngest) stratigraphic occurrence in this study. established biostratigraphic events are indicated by black text in figures, and new events are indicated in red text in figures. 9.1. palynology in total, the studied succession from the six wells covers 14 of the dinocyst zones defined by dybkjær & piasecki (2010), comprising the chiropteridium galea zone (early aquitanian) to the hystrichosphaeropsis obscura zone (late tortonian; figs 5, 6 and supplementary files s1–s6). in the 2/11–12s core, eight dinocyst zones were recorded, spanning the early burdigalian sumatradinium hamulatum zone to the serravallian gramocysta verricula zone (fig. 6). the studied succession in the partly cored 2/8–g10a well starts somewhat lower, possibly in the early aquitanian chiropteridium galea zone (fig. 5). the palynostratigraphy of this lower part is not well defined, and the presence of the lower two zones, the c. galea zone and the homotryblium spp. zone, are questionable as the index taxa occur in very low numbers and may be reworked. however, the occurrence of microfossil zone nsb9 (king 1989) around this level supports an early https://doi.org/10.34194/5k9dv133 http://www.geusbulletin.org/ sheldon et al. 2025: geus bulletin 59. 8381. https://doi.org/10.34194/5k9dv133 12 of 30 geusbulletin.org aquitanian age for this interval. the succession up to the unipontidinium aquaeductum zone seems to be complete. above the u. aquaeductum zone, three dinocyst zones, the achomosphaera andalousiense zone, gramocysta verricula zone and amiculosphaera umbracula zone, are missing corresponding to the upper part of the serravallian and the lower part of the tortonian succession. the u. aquaeductum zone is overlain by an interval referred to as the late tortonian hystrichosphaeropsis obscura zone. above the h. obscura zone is an interval referred to as the early serravallian a. andalousiense zone, indicating a repeated section possibly due to faulting. 9.2. dinocyst events in addition to the fos and los defining zonal boundaries, dybkjær & piasecki (2010, their figs. 6, 7) also presented selected additional stratigraphically useful events in their zonation. these events were also found in this study, including the base of ectosphaeropsis burdigalensis, the top of thalassiphora rota, the base of sumatradinium hamulatum, the top of exochosphaeridium insigne, the base of cerebrocysta poulsenii, the top of cousteaudinium aubryae, the base of palaeocystodinium miocaenicum, the top of unipontidinium aquaeductum, the occurrence of cannosphaeropsis passio and the top of palaeocystodinium miocaenicum. this study has revealed some differences when comparing the positions of these events with those of dybkjær & piasecki (2010). firstly, the top of exochosphaeridium insigne is found in both the 2/11–12s core and the 2/8g10a core above the base of cousteaudinium aubryae and thus within the c. aubryae zone. in dybkjær & piasecki (2010), this event was located below the base of c. aubryae, within the e. insigne zone. next, the top of cousteaudinium aubryae is located in both the 2/11–12s core and the 2/8–g10a core, somewhat above the fo of labyrinthodinium truncatum, and is thus within the l. truncatum zone. these two events were erroneously shown to occur at the same level (dybkjær & piasecki 2010, their figs. 6, 7), while in the description of the l. truncatum zone, it was stated that the lo of c. aubryae occurs within that zone. the data from this study thus support the latter. acccording to dybkjær & piasecki (2010), the top of cerebrocysta poulsenii is found at the top of the achomosphaera andalousiense zone. in this study, in the 2/11–12s core, two specimens of c. poulsenii were found in the same sample as the base of gramocysta verricula, the latter defining the base of the g. verricula zone. unfortunately, that sample was the highest sample analysed, so it is uncertain if the top of c. poulsenii continues further up in the g. verricula zone in the study area. lastly, the lo of cleistosphaeridium placacanthum was erroneously located at two different levels in the zonation of dybkjær & piasecki (2010): at the top of the achomosphaera andalousiense zone (p. 18) and coinciding with the upper boundary of the a. umbraculum zone (p. 19) and the lower boundary of the h. obscura zone (p. 21). the location of the top of cleistosphaeridium placacanthum is clearly not a good stratigraphic marker. this study indicates a gradual decrease in abundance of this species, resulting in an indistinct and poorly defined top. in the 2/11–12s core, the top of c. placacanthum was found in the upper part of the u. aquaeductum zone. in the 2/8–g10a core, it is present in the interval referred to as the a. andalousiense zone. in the 2/8–8 well, this species was found consistently in the upper part of u. aquaeductum zone and sporadically at least up to and within the uppermost analysed sample at the base of the g. verricula zone. in the 2/8–n4 well, the top occurrence of this species was found in an interval which either belongs to the upper u. aquaeductum zone or the lower a. andalousiense zone, and in the 2/11–1 well, the top of c. placacanthum was found in the upper part of the u. aquaeductum zone. however, the succession above the base of the g. verricula zone was not included in the cored sections in this study, and so, our data set does not provide a well-documented location for the lo of this species. new events for the danish and southern norwegian north sea area with potential stratigraphic use found in this study are given as follows (see figs 5, 6, 7, 8 and supplementary files s1–s13): • the lo of ectosphaeropsis burdigalensis in either the  homotryblium spp. zone or in the caligodinium amiculum zone in the 2/8-g10a well. • the lo of membranilarnacea cf. picena group in the upper caligodinium amiculum zone in the 2/11-1 well. • the lo of leptodinium italicum in the upper sumatradinium hamulatum zone in the 2/11-1 well. • the lo of dinopterygidium cladoides in the upper s. hamulatum zone in the 2/8g10a core (and also in the 2/8–8 and 2/11–1 wells). • the lo of hystrichokolpoma cinctum in the upper sumatradinium hamulatum zone in the 2/11-1 well. • the fo of invertocysta tabulata in the lower part of the labyrinthodinium truncatum zone in all wells. • the fo of habibacysta tectata in the lower part of the unipontidinium aquaeductum zone in the 2/11–12s and 2/8–g10a cored wells and also in the 2/8–8 well. • the fo of palaeocystodinium powellense in the unipontidinium aquaeductum zone in the 2/11-1 well. • the fo of operculodinium eirikianum in the middle part of the unipontidinium aquaeductum zone in the 2/11– 12s, 2/8–8 and 2/11–1 wells and in the middle to ?upper https://doi.org/10.34194/5k9dv133 http://www.geusbulletin.org/ sheldon et al. 2025: geus bulletin 59. 8381. https://doi.org/10.34194/5k9dv133 13 of 30 geusbulletin.org part of the unipontidinium aquaeductum zone in the 2/8–g10a well (the upper part of the zone has probably been removed by erosion or faulting, see fig. 5). • the fo of common habibacysta tectata in the middle part of the unipontidinium aquaeductum zone in the 2/11–12s, 2/8–8 and 2/11–1 wells and in the middle to ?upper part of the unipontidinium aqueaductum zone in the 2/8–g10a well. • the lo of sumatradinium hamulatum in the upper part of the unipontidinium aquaeductum zone in the 2/11–12s well, and in the achomosphaera andalousiense zone in the 2/8–n4, 2/8–8 and 2/11–1 wells. • the fo of common impagidinium sp. a wrenn & kokinos 1996 in the uppermost part of the unipontidinium aquaeductum zone in the 2/11–12s and 2/8– g10a wells. • the fo of invertocysta lacrymosa in the basal part of the achomosphaera andalousiense zone in the 2/11–12s well. • the fo of operculodinium tegillatum in the hystrichosphaeropsis obscura zone in the 2/8-g10a well. • the lo of impagidinium sp. a wrenn & kokinos 1996 (common) in the achomosphaera andalousiense zone in the 2/8-g10a well. established and new dinocyst events noted in this study are illustrated in figs 5, 6 and supplementary files s1– s6. the most useful dinocyst events for correlation on a local level are marked using purple correlation lines in supplementary files s7–s12 and combined in a correlation figure (fig. 8 and supplementary file s13). 9.3. microfossils application of the north sea microfossil zonations, nsb (king 1989) and ns (king 2016), referred the combined studied interval of the six wells to aquitanian to early tortonian zones nsb9 (ns34) to nsb12c (ns38–39; figs 4–6, supplementary files s1–s12). the calcareous benthic foraminifera that mark the top of nsb9 (ns34; lo of plectofrondicularia seminuda), top of nsb10 (ns35; lo of uvigerina teniupustulata), top of nsb11 (ns36b; lo of asterigerina guerichi staeschei) and top of nsb12 (ns39; lo of uvigerina kingii), along with other marker planktonic microfossils from king (1989, 2016) are applied successfully in this study. the 2/11–12s core spans burdigalian zone nsb10 (ns35a) to serravallian subzone nsb12c (ns39). the 2/8–g10a core covers aquitanian zone nsb9 (ns34) langhian to subzone nsb12a (ns36c). while fos and los of planktonic and calcareous benthic foraminifera are useful for biostratigraphic correlation in the valhall–hod area, abundance variations in other microfossil groups, such as agglutinating foraminfera, diatoms, radiolaria and sponge spicules, are also potentially useful for correlation. in addition, variations in the siliceous versus calcareous microfossil components are of interest regarding our understanding of the diatomite reservoir architecture and palaeoenvironment and form the basis for ongoing detailed studies. 9.4. microfossil events in addition to the fos and los of microfossils defining zone boundaries, king (1989, 2016) also presented selected additional stratigraphically useful events. this study reveals some local observations and differences in the relative positions of some of these events (see figs. 5, 6, 8 and supplementary files s1–s13) as follows: • the fo of globorotalia praescitula below the fo of uvigerina tenuipustulata in the 2/11–12s and 2/8–g10a wells. in king (1983), these two events are in reverse order, although their ranges are marked as uncertain. • the fo of globorotalia zealandica above the fo of uvigerina tenuipustulata in the 2/11–12s and 2/8– g10a wells. in king (1983), these two events coincide, although their ranges are noted as uncertain. • the coinciding los of globorotalia zealandica and globorotalia praescitula just below the top of zone nsb10 (ns35b-c) in the 2/11–12s and 2/8–g10a wells. in king (1983, 2016), the lo of globorotalia zealandica occurs slightly earlier in nsb10 than the lo of globorotalia praescitula. • the lo of trilobatus trilobus is noted within zone nsb11 (ns36a-b) in this study. king (1983) places its lo in the lower part of subzone nsb12a. • the lo of loxostomum sinuosum is seen at the top of zone nsb11 (ns36a-b) in the 2/8–g10a, 2/11–12s, 2/8–n4 and 2/11–1 wells. the lo of loxostomum sinuosum is found in the lower part of subzone nsb10 (laursen & kristoffersen 1999) onshore denmark, at the top of nsb11 in king (1989) and at the top of nsb9 (ns34) in king (2016). • the fo of uvigerina kingi marks the base of nsb12c (ns38–39) in this study in the absence of elphidium antoninum (king 1989). in king (1989), the fo of uvigerina kingi is within subzone nsb12c with uncertainty. we also identified new microfossil events not noted in the zonations of king (1983, 1989, 2016). these events have potential stratigraphic use and are found in the cored intervals in this study and also in some of the non-cored wells (see figs. 5, 6, 8 and supplementary files s1–s13, where they are noted in red). these events are as follows: • the lo of globigerinelloides obliquus in the upper part of nsb10 (ns35b-c) below the los of alabamina wolterstorffi and trifarina bradyi and above the fo of uvigerhttps://doi.org/10.34194/5k9dv133 http://www.geusbulletin.org/ sheldon et al. 2025: geus bulletin 59. 8381. https://doi.org/10.34194/5k9dv133 14 of 30 geusbulletin.org ina tenuipustulata in the 2/8–g10a and 2/11–12s wells. • the lo of ceratobulimina contraria just below (2/8–g10a) or close to (2/11–12s, 2/8–v6, 2/11–1) the lo of uvigerina tenuipustulata in upper nsb10 (ns35b-c). onshore denmark, in a shallower setting, the lo of ceratobulimina contraria is a younger event marking the top of nsb12a (laursen & kristoffersen 1999). • the los of alabamina wolterstorffi and trifarina bradyi in the upper part of nsb10 (ns35b-c) below the los of globorotalia zealandica and globorotalia praescitula in the 2/8–g10a and 2/11–12s wells. the lo of trifirina bradyi is found in the upper pliocene succession in the north sea (king 1989). its consistent lo towards the top of nsb10 in the study area (also in the 2/8–n4, 2/8–8 and 2/11–1 wells) may be useful for local correlation. • the los of ciperoella ciperoensis and paragloborotalia nana close to the base of nsb11 (ns36a-b) in the 2/8– g10a, 2/11–12s, 2/8–n4 and 2/8–v6 wells. the lo of ciperoella ciperoensis was noted in the lower miocene succession of the ekofisk field (eidvin et al. 1999). • the lo of globoturborotalita eolabiacrassata in the lower part of nsb11 (ns36a-b) in the 2/8–g10a and 2/11–12s wells. • the lo of lenticulina peregrina at or just below the top of nsb11 (ns36a-b) in the 2/8–g10a, 2/11–12s and 2/8–n4 wells. • the lo of cancris subconicus at the top of nsb11 (ns36a-b) in the 2/8–g10a and 2/11–12s wells. two additional events are identified that may have correlation potential (also marked in red in supplementary files s1–s12). these include the lo of nonion cf. granosum towards the top of nsb10 (ns35b-c) in the 2/8–g10a, 2/11–12s, 2/8–n4, 2/8–8 and 2/8–v6 wells and the los of ciperoella anguliofficionalis and trilobatus quadrilobatus within subones nsb12a-b (ns36c-38) in the 2/8–g10a, 2/11–12s, 2/8–n4 and 2/8–8 wells. established and new microfossil events noted in this study are illustrated in figs. 5, 6 and supplementary files s1–s6. the most useful microfossil events for correlation on a local level are marked using blue correlation lines in supplementary files s7–s12 and combined into a correlation figure (fig. 8 and supplementary file s13). 9.5. calcareous nannofossils approximately ten samples per well were analysed for calcareous nannofossils using basic presence–absence observations, which enabled a broad biostratigraphic breakdown into fairly long-ranging nannofossil zones, overall spanning early to middle miocene zones nn3– nn6 (martini 1971). calcareous nannofossil ranges are based on observations of young (1998), de kaenel et al. (2017), bergen et  al. (2017) and boesiger et  al. (2017). reworked nannofossils from the upper cretaceous and palaeogene are present in all wells, and caved material was occasionally noted in the non-cored wells. nannofossil events in figs. 5, 6 and supplementary files s1–s12 are in black text and are not included in the correlation in fig. 8 and supplementary file s13 due to large sample spacing and the basic counting method applied. the relative position of early to middle miocene calcareous nannofossil events from the valhall–hod area is shown in fig. 7. while these events are not necessarily new, they may be useful for local and regional correlation and biostratigraphy. the oldest nannofossil zone identified is zone nn3 in the 2/8–8 well, recognised due to the co-occurrence of helicosphaera bipuncta, sphenolithus apoxis, sphenolithus puniceus and helicosphaera scissura. joint zones nn3–4 are noted in the 2/8–g10a, 2/11–12s and 2/8–v6 wells based on an assemblage containing elements restricted to either zone. for example, sphenolithus apoxis and sphenolithus conicus (los in nn3), discoaster emblematicus (fo in nn3) and discoaster caulifloris, discoaster petaliformis and sphenolithus heteromorphus (fos in nn4). zone nn4 is identified in all wells and is characterised by the co-occurrence of helicosphaera ampliaperta, variably with helicosphaera waltrans, sphenolithus puniceus, sphenolithus heteromorphus, helicosphaera scissura, discoaster emblematicus, discoaster caulifloris and sphenolithus abies. joint zone nn4–5 is recorded in the 2/8– g10a well, based on the co-occurrence of helicosphaera scissura, discoaster caulifloris, reticulofenestra pseudoumbilicus and discoaster discissus. the upper parts of the 2/8–n4, 2/8–8, 2/8–v6 and 2/11–1 wells are assigned a wide nn4–6 range based on the presence of cyclicargolithus bukryi, cyclicargolithus floridanus and helicosphaera vedderi whose upper range is no younger than zone nn6. 9.6. calcareous nannofossil events nannofossil events with potential correlative use and other possible useful nannofossil biostratigraphic occurrences are seen in figs. 5, 6, 7 and supplementary files s1–12. they are described here in stratigraphic order in relation to nsb/ns microfossil zones (king 1989, 2016), and their relative positioning is shown in fig. 7 against the nannofossil zonation of martini (1971). low resolution sample spacing and potential caving can result in depressed fos in the non-cored wells. the events are indicated by the: • fo of coccolithus pelagicus in all wells at the base of the studied sections. • fo of cyclicargolithus floridanus in all wells at or near https://doi.org/10.34194/5k9dv133 http://www.geusbulletin.org/ sheldon et al. 2025: geus bulletin 59. 8381. https://doi.org/10.34194/5k9dv133 15 of 30 geusbulletin.org the base of the studied sections. • fo of helicosphaera carteri at the base of the studied sections in the 2/8–g10a, 2/11–12s, 2/8–n4, 2/8–8 and 2/11–1 wells. • fo of helicosphaera ampliaperta at the base of the studied sections in the 2/11–12s, 2/8–n4, 2/8–8 and 2/11–1 wells and slightly higher, in the upper part of nsb10 (near the base of ns35b) in the 2/8–g10a and 2/8–v6 wells. • fo of discoaster emblematicus at the base of the studied sections in the 2/8–g10a, 2/8–n4 and 2/8–8 wells and higher, in the upper part of nsb10 (ns35b-c) in the 2/11–12s and 2/8–v6 wells. • fo of reticulofenestra pseudoumbilicus in the 2/11– 12s, 2/8–v6 and 2/11–1 wells towards the middle of nsb10 (ns35). • fo of discoaster caulifloris in the upper part of nsb10 (near the base of ns35b) in the 2/8–g10a, 2/11–12s and 2/11–1 wells and slightly higher in the 2/8–v6 well. • fo of discoaster exilis in the 2/8–n4, 2/8–8 and 2/11–1 wells towards the middle of nsb10 (near the base of ns35b). • lo of discoaster emblematicus close to the boundary of nsb10 (ns35b-c) and nsb11 (ns36a) in the 2/11– 12s, 2/8–8 and 2/8–v6 wells, and slightly lower in the 2/g10a and 2/8–n4 wells. • fo of helicosphaera bipuncta in the 2/8–g10a, 2/11– 12s and 2/8–v6 wells towards the base of nsb11 (ns36a-b) and lower, perhaps due to caving in the 2/11–1 well. • lo of helicosphaera ampliaperta in nsb10 (ns35) in the 2/8–g10a, 2/8–n4, 2/8–8 and 2/11–1 wells and higher, within nsb12a (ns36c) in the 2/11–12s and 2/8–v6 wells. the discrepancy in age may be due to wide nannofossil sample spacing and the thin succession represented by nsb11. • lo of discoaster exilis in nsb12a-b (ns36c-38) in the 2/8–n4, 2/8–8 and 2/11–1 wells. • lo of discoaster caulifloris within nsb11 (ns36a-b) in the 2/8–g10a and 2/11–12s wells and slightly lower, in upper nsb10 (ns35), in 2/11–1. • lo of helicosphaera bipuncta in nsb12 (ns36–38) in the 2/11–12s and 2/11–1 wells, and lower in nsb11 (ns36a-b) in the 2/8–g10a and 2/8–v6 wells. • lo of cyclicargolithus floridanus at or near the top of the studied sections in all wells. • lo of helicosphaera carteri at or near the top of the studied sections in all wells. • lo of coccolithus pelagicus in all wells at the top of the studied sections. a more extensive study of nannofossils of the calcareous intervals was unfortunately beyond the scope of this study but would be an interesting exercise to carry out in the future. 9.7. diatoms sufficiently well-preserved diatoms recorded in the cored and non-cored wells in this study allowed the application of the early and middle miocene biostratigraphy of schrader and fenner (1976) from the norwegian sea. in this study, the early miocene rhizosolenia norwegica–coscinodiscus vigilans–nitzschia maleinterpretaria joint partial range zone (prz) to the middle miocene denticulopsis hyalina prz are recognised (figs. 5, 6, 8, supplementary files s1–s13). zone definitions of schrader & fenner (1976) and observed assemblages in this study are described here. apart from in the 2/8–g10a core, several of the diatom species that define the tops and bases of the przs of schrader & fenner (1976) are not consistently recorded in this study. however, the diatom assemblages described in schrader & fenner (1976) in a particular prz are similar to the assemblages found here and can provide an alternative method of identifying that specific prz when the markers for prz tops and bases are not seen. new diatom events are recognised in this study within the identified przs in several wells. therefore, these new events correlate across the valhall and hod areas and may potentially be of correlative fig. 7 overview of the relative position of early to middle miocene calcareous nannofossil events from the valhall–hod area. nannofossil zonation of martini (1971). age (ma) 11 12 13 14 15 16 17 18 19 20 21 22 23 epoch stage nannofossils zone m io ce ne se rr av al lia n la ng hi an bu rd ig al ia n a qu it an ia n to rt on ia n nn8 nn7 nn6 nn5 nn4 nn3 nn2 nn1 c. pelagicus, h. carteri, c. �oridanus, h. ampliaperta, d. emblematicus d. exilis, d. cauli�oris, r. pseudoumbilicus c. pelagicus, h. carteri, c. �oridanus h. bipuncta d. cauli�oris d. exilis h. ampliaperta d. emblematicus h. bipuncta https://doi.org/10.34194/5k9dv133 http://www.geusbulletin.org/ sheldon et al. 2025: geus bulletin 59. 8381. https://doi.org/10.34194/5k9dv133 16 of 30 geusbulletin.org value further afield. the chronostratigraphic age of the diatom zonations in this study (fig. 4) is linked to the dinocyst stratigraphy of dybkjær & piasecki (2010) in the description that follows. 9.8. diatom zonation rhizosolenia norwegica–coscinodiscus vigilans–nitzschia maleinterpretaria joint prz (schrader & fenner 1976) definition. the base of the rhizosolenia norwegica prz is based on the fo of stictodiscus aff. kittonianus and macrora stella. the top of the nitzschia maleinterpretaria prz is based on the los of dimerogramma fossile, sceptroneis ossiformis, thalassionema harosakiensis and thalassiosira spinosa var. aspinosa and the fo of raphoneis margaritalimbata. age. early miocene (burdigalian). from the lower part of the upper sumatradinium hamulatum zone to the lower part of the cousteaudinium aubryae dinocyst zone (dybkjær & piasecki 2010) in the 2/8–g10a cored well. diatom assemblage (this study). includes stictodiscus aff. kittonianus, raphoneis margaritalimbata, opephora gemmata, raphidodiscus marylandicus, pseudodimerogramma elongata, thalassiosira fraga, rocella gelida, rhizosolenia hebetata, sceptroneis ossiformis, actinocyclus ehrenbergii, dimerogramma fossile, actinocyclus tenellus, stephanopyxis horridus, diploneis smithii, thalassionema spp. (abundant in the uppermost part), synedra jouseana, raphoneis amphiceros, paralia sulcata (abundant), chaetoceros spp. (abundant), pseudopodosira spp. (abundant, including pseudoporosira westii) and actinoptychus senarius. observations. it was not possible to subdivide these three przs in this study. in the 2/8–g10a core, the base of this combined interval was based on the fo of stictodiscus aff. kittonianus and the top by the fo of raphoneis margaritalimbata. the fo of stictodiscus aff. kittonianus is found with the fo of opephora gemmata in the upper sumatradinium hamulatum dinocyst zone in the 2/8–g10a core. the fo of opephora gemmata is found at a similar stratigraphic level in the 2/8–8, 2/8– v6 and 2/11–1 wells and is here used as an additional diatom event for the identification of the base of the r. norwegica–c. vigilans–n. maleinterpretaria joint prz. the top of this joint prz is marked by the fo of raphoneis margaritalimbata in the 2/8–g10a, 2/8–8 and 2/11–1 wells (figs. 5, 8, supplementary files s1, s4, s6, s7, s10, s12, s13), close to the fo of the dinocyst cousteaudinium aubryae. the top of this prz in the 2/8–v6 well (supplementary files s5, s11) is comparatively high, but due to a lack of other marker diatoms is based on the fo  of common to abundant stephanopyxis horridus in the overlying sample, an event which is characteristic of the overlying thalassiosira fraga prz. the r. norwegica–c. vigilans–n. maleinterpretaria joint prz is only identified in the 2/8–g10a, 2/8–8, 2/8–v6 and 2/11–1 wells. thalassiosira fraga prz (schrader & fenner 1976) definition. the base of the thalassiosira fraga prz is defined on the los of dimerogramma fossile, sceptroneis ossiformis, thalassionema harosakiensis and thalassiosira spinosa var. aspinosa, and the fo of raphoneis margaritalimbata. the top is based on the fos of coscinodiscus lewisianus, cymatosira biharensis, dimerogramma aff. dubium and hemiaulus malleus and the lo of thalassiosira fraga. age. early to middle miocene (burdigalian to early langhian). from the lower part of the cousteaudinium aubryae zone to the lower part of the labyrinthodinium truncatum dinocyst zone (dybkjær & piasecki 2010) in the 2/8–g10a cored well. diatom assemblage (this study). includes chaetoceros spp. (abundant), thalassiosira fraga, cymatosira biharensis, raphoneis margaritalimbata, thalassionema nitzschiodes (abundant), paralia sulcata (abundant), pseudopodosira spp. (common, including pseudopodosira westii), actinocyclus ingens (rare), actinocyclus ehrenbergii, raphoneis robustata, raphoneis angulata, raphidodiscus marylandicus, pseudodimerogramma elongata, stephanopyxis horridus, diploneis smithii, stictodiscus aff. kittonianus, synedra jouseana, stephanopyxis grunowii, rhizosolenia hebetata, rhizosolenia miocenica (rare), raphoneis amphiceros, pterotheca reticulata, opephora gemmata, cestodiscus peplum and actinoptychus senarius. observations. in the 2/8–g10a, 2/8–8 and 2/11–1 wells, the base of this prz is based on the fo of raphoneis margaritalimbata (fig. 5, supplementary files s1, s4, s6, s7, s10, s12) close to the fo of the dinocyst cousteaudinium aubryae. in the 2/8–g10a well, the top of this prz is marked by the fo of cymatosira biharensis coinciding with the lo of thalassiosira fraga (fig. 5, supplementary files s1, s7) and in the 2/8–8 well (supplementary files s4, s10) by the fo of cymatosira biharensis, in the labyrinthodinium truncatum dinocyst zone, below the lo of the dinocyst cousteaudinium aubryae. in the absence of the aforementioned established markers, alternative diatom events recognised in this study in the thalassiosira fraga prz are the fo of common to abundant stephanopyxis horridus close to the base of the labyrinthodinium truncatum dinocyst zone in the 2/8– g10a, 2/11–12s, 2/8–8, 2/8–v6 and 2/11–1 wells. the fo of common to abundant thalassionema nitzschiodes https://doi.org/10.34194/5k9dv133 http://www.geusbulletin.org/ sheldon et al. 2025: geus bulletin 59. 8381. https://doi.org/10.34194/5k9dv133 17 of 30 geusbulletin.org is also mainly noted within the thalassiosira fraga prz in this study, but variably within the cousteaudinium aubryae and labyrinthodinium truncatum dinocyst zones, suggesting that this event is diachronous. the lo of raphoneis margaritalimbata is noted towards the top of the cousteaudinium aubryae dinocyst zone in the 2/8– g10a, 2/8–8 and 2/11–1 wells, and the fo of raphoneis robustata is noted at the same level in the 2/8–g10a and 2/11–1 wells. both events are potentially useful. the t. fraga prz is identified in the 2/8–g10a, 2/11–12s, 2/8–8, 2/8–v6 and 2/11–1 wells (figs. 5, 6, supplementary files s1–s2, s4–s8, s10–s12). rhizosolenia bulbosa prz (schrader & fenner 1976) definition. the base of the rhizosolenia bulbosa prz is defined on the fos of coscinodiscus lewisianus, cymatosira biharensis, dimerogramma aff. dubium and hemiaulus malleus and the lo of thalassiosira fraga, and the top on the fos of denticulopsis hustedtii, denticulopsis norwegica, coscinodiscus endoi and rhizosolenia miocenica. age. middle miocene (langhian). from the upper part of the labyrinthodinium truncatum zone to the lower part of the unipontidinium aquaeductum dinocyst zone (dybkjær & piasecki 2010) in the 2/8–g10a cored well. diatom assemblage (this study). includes cymatosira biharensis, cestodiscus peplum, raphoneis robustata, raphoneis angulata, pseudodimerogramma elongata, actinocyclus tenellus, actinocyclus ingens (rare), diploneis smithii, actinoptychus splendens, synedra jouseana, stephanopyxis horridus (abundant), stictodiscus aff. kittonianus, thalassionema nitzschiodes (abundant), chaetoceros spp. (abundant), mediaria splendida (rare), paralia sulcata (common), pseudopodosira spp. (common, including pseudopodosira westii), raphidodiscus marylandicus, stephanopyxis turris, stephanopyxis grunowii, rhizosolenia miocenica (rare), rhizosolenia hebetata, raphoneis amphiceros, synedra spp., pterotheca reticulata, opephora gemmata, denticulopsis spp. (rare) and actinoptychus senarius. observations. in the 2/8–g10a core, the base of this prz was based on the co-occurrence of the fo of cymatosira biharensis and the lo of thalassiosira fraga, and in the 2/8–8 well by the fo of cymatosira biharensis. in the 2/8–g10a, 2/11–12s, 2/8–n4, 2/8–8 and 2/8–v6 wells, the top of this prz is based on the fo of denticulopsis hustedtii. the diatoms that define the base of the rhizosolenia bulbosa prz were only recorded in the 2/8– g10a and 2/8–8 wells. in the absence of these marker diatoms in the other wells, alternative diatom events are recognised in this study in the rhizosolenia bulbosa prz, which are potentially useful for correlation. an influx of synedra spp. is noted in all wells towards the base of the rhizosolenia bulbosa prz, and the fo of consistent stephanopyxis grunowii is also noted towards the base of this zone in the 2/8–g10a, 2/8–8 and 2/8–v6 wells. the lo of stephanopyxis horridus (common to abundant) is noted in all wells towards the middle of this prz (figs. 5, 6, 8, supplementary files s1–13). the fo of denticulopsis hustedtii coincides in all wells with the fo of common to abundant denticulopsis spp. (including denticulopsis hyalina). this may be an easier event to recognise to mark the top of the rhizosolenia bulbosa prz. the rhizosolenia bulbosa prz is identified in all wells (figs. 5, 6, supplementary files s1–s13). denticulopsis hyalina prz (schrader & fenner 1976) definition. the base of the denticulopsis hyalina prz is defined on the fos of denticulopsis hustedtii, denticulopsis norwegica, coscinodiscus endoi and rhizosolenia miocenica. the top is based on the los of hemiaulus malleus, hemiaulus malleolus and pseudodimerogramma elongata and the fos of the coscinodiscus plicatus group and rouxia californica. age. middle miocene (late langhian). unipontidinium aquaeductum dinocyst zone (dybkjær & piasecki 2010) in the 2/8–g10a well. diatom assemblage (this study). includes denticulopsis hyalina, denticulopsis hustedtii, pseudodimerogramma elongata, mediaria splendida (rare), raphoneis wicomicoensis, actinocyclus tenellus, rhizosolenia miocenica (rare), rhizosolenia hebetata, cymatosira biharensis (rare), opephora gemmata, stephanopyxis horridus, diploneis smithii, actinoptychus splendens (rare), actinoptychus heliopelta (rare), actinocyclus ingens (common), stephanopyxis turris, stictodiscus aff. kittonianus, synedra jouseana (abundant), thalassionema nitzschiodes (abundant), stephanopyxis grunowii, raphoneis amphiceros, pterotheca reticulata, paralia sulcata (common), chaetoceros spp. (super-abundant) and actinoptychus senarius. pseudopodosira spp. (common, including pseudopodosira westii) is present throughout this prz and becomes abundant towards the top of the studied section. observations. in the 2/8–g10a, 2/11–12s, 2/8–n4, 2/8–8 and 2/8–v6 wells, the base of this prz was based on the fo of denticulopsis hustedtii. this event is coincident with the fo of common denticulopsis spp. (including denticulopsis hyalina, denticulopsis nicobarica (rare), denticulopsis punctata (rare), denticulopsis kanayae and denticulopsis lauta) on which the base of this prz is based in the 2/11–1 well. a series of local lo events occur within the denticulopsis hyalina prz: the lo of thalassionema nitzschiodes (abundant), cymatosira biharensis, denticulopsis https://doi.org/10.34194/5k9dv133 http://www.geusbulletin.org/ sheldon et al. 2025: geus bulletin 59. 8381. https://doi.org/10.34194/5k9dv133 18 of 30 geusbulletin.org hyalina, actinocyclus ingens, diploneis smithii, stictodiscus aff. kittonianus, synedra spp. and common denticulopsis spp. the top of the denticulopsis hyalina prz in this study is based on the lo of high abundance and diversity diatom assemblages without evidence of species from the overlying coscinodiscus plicatus group prz and approximates to the upper part of the unipontidinium aquaeductum dinocyst zone. the denticulopsis hyalina prz is identified in all wells (figs. 5, 6, 8, supplementary files s1–s13). 9.9. diatom events and observations from this study early to middle miocene diatom events found in the norwegian and iceland seas (schrader & fenner 1976; koç & scherer 1996) used in this study are presented and discussed here and noted in black in figs. 5, 6 and supplementary files s1–s12. the events are correlated with the dinocyst zonation of dybkjær and piasecki (2010; see figs. 5, 6, 8 and supplementary files s1–s13). selected diatom occurrences or events that were found in the norwegian sea and are noted sporadically in this study are found on figs. 5, 6, and supplementary files s1–s12 and are also noted in black on the figures. boundary-defining events of schrader and fenner’s (1976) zonation are discussed here as follows: • the fo of stictodiscus aff. kittonianus defines the base of the rhizosolenia norwegica prz (schrader & fenner 1976) and is only noted in the 2/8–g10a well at the base of the burdigalian (upper part of the sumatradinium hamulatum dinocyst zone). • the fo of raphoneis margaritalimbata defines the base of the thalassiosira fraga prz (schrader & fenner 1976) and is found in the 2/8–g10a, 2/8–8 and 2/11–1 wells close to the fo of the cousteaudinium aubryae dinocyst in the middle burdigalian stage (fig. 5, supplementary files s1, s4, s6, s7, s10 and s12). the fo of raphoneis margaritalimbata is found in the synedra pulchella interval (dated as late middle miocene by koç & scherer 1996) in the iceland sea. • the lo of thalassiosira fraga marks the top of the thalassiosira fraga prz and the base of the rhizosolenia bulbosa prz. this early miocene event (schrader & fenner 1976) was only noted in the 2/8–g10a well in the middle of the early langhian labyrinthodinium truncatum dinocyst zone. • the fo of cymatosira biharensis is recognised in the 2/8–g10a and 2/8–8 wells in the middle of the early langhian labyrinthodinium truncatum dinocyst zone. its fo defines the base of the rhizosolenia bulbosa prz in the norwegian sea (schrader & fenner 1976), which they interpreted as a late early miocene event. in the iceland sea, its fo is noted towards the base of the proboscia praebarboi interval, dated as early late miocene by koç & scherer (1996). • the fo of actinocylus ingens (2/11–12s, 2/8–n4) or the fo of consistent actinocylus ingens (2/8–g10a, 2/8–8, 2/8–v6, 2/11–1) is seen in all wells in the middle to upper part of the labyrinthodinium truncatum dinocyst zone, close to the lo of cousteaudinium aubryae. in the norwegian sea, the fo of actinocylus ingens is found in the rhizosolenia bulbosa prz of schrader & fenner (1976), which they interpreted to be late early miocene in age. in the iceland sea, its fo is found close to the base of the actinocylus ingens interval (dated as late middle miocene by koç & scherer 1996). • in the norwegian sea, the fo of rhizosolenia miocenica defines the base of the denticulopsis hyalina prz (and the top of the underlying rhizosolenia bulbosa prz), schrader & fenner (1976). the fo of rhizosolenia miocenica is seen in the 2/8–g10a, 2/8–n4, 2/8–v6 and 2/11–1 wells in the lower part of the labyrinthodinium truncatum dinocyst zone, in the upper part of the thalassiosira fraga and lowermost rhizosolenia bulbosa prz’s. while noting that this event defines a zonal boundary in schrader & fenner (1976), in this study, other events are deemed to be more useful. additional diatom events (fos, los) noted in schrader & fenner (1976), and observations from this study with potential stratigraphic use in the north sea, are presented here (see figs. 5, 6, 8 and supplementary files s1–s13; marked on the figures in red) as follows: • the fo of high abundance and diversity diatom assemblages is found in the upper part of the sumatradinium hamulatum dinocyst zone in the 2/8–g10a, 2/8–8, 2/8–v6 and 2/11–1 wells. in the 2/11–12s well, the event is found at the base of the cousteaudinium aubryae dinocyst zone, and in the 2/8–n4 well, it is found towards the base of the labyrinthodinium truncatum dinocyst zone, clearly displaying that this event is diachronous across the valhall–hod area. • the fo of opephora gemmata is found in the 2/8–g10a, 2/8–8, 2/8–v6 and 2/11–1 wells close to or at the fo of high abundance and diversity diatom assemblages, in the upper part of the sumatradinium hamulatum dinocyst zone. the fo of opephora gemmata is found at the base of the synedra jouseana prz (dated as early miocene by schrader & fenner 1976) in the norwegian sea. • the fo of common to abundant thalassionema nitzschiodes is noted in all wells, usually in the middle cousteaudinium aubryae dinocyst zone (wells 2/8–g10a, 2/8–8, 2/8–v6, 2/11–1) and occasionally in the lower labyrinthodinium truncatum zone (2/11–12s, 2/8–n4). thalassionema nitzschiodes is https://doi.org/10.34194/5k9dv133 http://www.geusbulletin.org/ sheldon et al. 2025: geus bulletin 59. 8381. https://doi.org/10.34194/5k9dv133 19 of 30 geusbulletin.org consistently abundant in the iceland sea from the top of the synedra pulchella interval to the top of the denticulopsis hustedtii interval, dated as late middle miocene to early late miocene by koç & scherer (1996). • the lo of raphoneis margaritalimbata is recognised in the 2/8–g10a, 2/8–8 and 2/11–1 wells and wells towards the top of the cousteaudinium aubryae dinocyst zone. in the iceland sea, this event occurs in the probiscia praeparboi interval, dated early late miocene according to koç & scherer (1996), and in the norwegian sea at the base of the cymatosira biharensis prz, dated late miocene by schrader & fenner (1976). • the fo of common to abundant stephanopyxis horridus is recognised in all wells, usually in the upper cousteaudinium aubryae dinocyst zone, occasionally in the lower labyrinthodinium truncatum dinocyst zone. in the norwegian sea, the base of stephanopyxis horridus is not well defined but it is present at least from the actinocyclus ingens prz, dated as middle miocene by schrader & fenner (1976). • the fo of raphoneis robustata is recognized in the 2/8–g10a, 2/8–n4 and 2/11–1 wells in the upper part of the cousteaudinium aubryae or lower labyrinthodinium truncatum dinocyst zones. • the fo of an influx of synedra spp. is noted in all wells at the base of the rhizosolenia bulbosa prz in this study, in the lower to mid labyrinthodinium truncatum dinocyst zone. • the lo of stephanopyxis horridus (common to abundant) is recognised in all wells in the labyrinthodinium truncatum dinocyst zone. in the norwegian sea, the lo of stephanopyxis horridus is found towards the base of the rhizosolenia miocenica prz, dated as early late miocene by schrader & fenner (1976). • the fo of consistent stephanopyxis grunowii is recorded in all wells in the labyrinthodinium truncatum dinocyst zone, close to the lo of stephanopyxis horridus (common to abundant). • the fo of denticulopsis hyalina is noted close to the boundary of the labyrinthodinium truncatum and unipontidinium aquaeductum dinocyst zones in the 2/8–g10a, 2/11–12s, 2/8–n4, 2/8–8 and 2/11–1 wells, just below or at the fo of common to abundant denticulopsis spp. in the norwegian sea, the fo of denticulopsis hyalina is close to the base of the denticulopsis hyalina prz, dated as early to middle miocene by schrader & fenner (1976). • the fo of common to abundant denticulopsis spp. is noted in all wells and is found towards the base of the unipontidinium aquaeductum dinocyst zone in the 2/11–12s, 2/8–g10a, 2/8–8 and 2/8–v6 wells and towards the top of the upper labyrinthodinium truncatum dinocyst zone in the 2/8–n4 and 2/11–1 wells. denticulopsis spp. is consistently common to abundant from the coscinodiscus norwegicus interval, dated as middle miocene by koç & scherer (1996) in the iceland sea and the denticulopsis hyalina prz, dated as early middle miocene by schrader & fenner (1976) in the norwegian sea. common denticulopsis spp. is noted in the upper langhian interval in the e-8x well, danish sector of the north sea. • the lo of common-abundant thalassionema nitzschiodes is recorded in the unipontidinium aquaeductum dinocyst zone in most wells, not far above the fo of common to abundant denticulopsis spp. in the iceland sea, this event is noted in the proboscia barboi interval, dated as late miocene by koç & scherer (1996). • the lo of diploneis smithii is recognised at or between the lo of common-abundant thalassionema nitzschiodes and the lo of high abundance and diversity diatom assemblages in the unipontidinium aquaeductum dinocyst zone in the 2/8–g10a, 2/11–12s, 2/8–v6 and 2/8–8 wells and at the top of the labyrinthodinium truncatum dinocyst zone in the 2/11–1 well. the ‘diploneis group’ is noted as rare occurrences in the high-latitude north atlantic ocean region from the early miocene to the early pliocene in baldauf (1982). • the lo of actinocyclus ingens is recognised in all wells in the upper part of the unipontidinium aquaeductum dinocyst zone. in this study, its lo is close to the lo of diploneis smithii. in the 2/8–n4 and 2/8–8 wells, its lo coincides with the lo of high abundance and diversity diatom assemblages. in the norwegian sea, the lo of actinocyclus ingens marks the top of the goniothecium tenue prz (schrader & fenner 1976), which is middle miocene in age according to barron (1985). in the iceland sea, its lo marks the top of the proboscia praebarboi interval, dated as early late miocene by koç & scherer (1996). • the lo of denticulopsis hyalina is recognised in the 2/8–g10a, 2/11–12s, 2/8–8, 2/11–1 and 2/8–n4 wells just below the lo of actinocyclus ingens and in most sections just below the lo of high abundance and diversity diatom assemblages. in the norwegian sea, the lo of denticulopsis hyalina is found at the top of the coscinodiscus plicatus prz, dated as middle miocene by schrader & fenner (1976). • the lo of common denticulopsis spp. (including denticulopsis lauta and denticulopsis hustedtii) is recorded in most wells in the upper part of the unipontidinium aquaeductum dinocyst zone. in the norwegian sea and iceland sea regions, the event is found in the denticulopsis hustedtii prz and interval, dated as late miocene by schrader & fenner (1976) and koç & https://doi.org/10.34194/5k9dv133 http://www.geusbulletin.org/ sheldon et al. 2025: geus bulletin 59. 8381. https://doi.org/10.34194/5k9dv133 20 of 30 geusbulletin.org scherer (1996). in the e-8x well in the danish sector of the north sea, the lo of denticulopsis spp. is noted in the late langhian stage. this event coincides with the lo of high abundance and diversity diatom assemblages in the 2/8–g10a, 2/11–12s, 2/8–8 and 2/11–1 wells and is slightly below this level in the 2/8–n4 and 2/8–v6 wells. • the lo of high abundance and diversity diatom assemblages (including opephora gemmata, stephanopyxis grunowii and synedra spp.) is recognised in most wells in this study in the upper part of the unipontidinium aquaeductum dinocyst zone. this event coincides with the top of the hodde formation in the 2/8–g10a, 2/11–12s, 2/8–8 wells and the top of the nora formation in the 2/8–n4, 2/8–v6 and 2/11–1 wells. diatom events noted in this study are illustrated in figs.  5, 6 and supplementary files s1–s6. the most useful diatom events for correlation on a local level are marked using green correlation lines in supplementary files s7–s12 and combined in the correlation figure (fig. 8 and supplementary file s13). 9.10. silicoflagellates silicoflagellates commonly only make up about 2–3% of the biogenic part of siliceous sediments (mccartney et al. 2011). as a silicoflagellate zonation for the north sea does not exist, the early and middle miocene norwegian sea silicoflagellate biostratigraphy of locker & martini (1989) was applied, with some success, to the 2/8–g10a well and with limited success to the 2/8–n4, 2/8–8 and 2/11–1 wells (fig. 5, and supplementary files s1, s3, s4, s6, s7, s9, s10, s12). the 2/11–12s and 2/8–v6 wells did not yield assemblages adequate for a silicoflagellate biostratigraphic breakdown. two silicoflagellate zones are identified in this study: (1) the lower corbisema triacantha zone and (2) the upper corbisema triacantha zone. as this is the first time that silicoflagellate biostratigraphy has been applied to north sea lower-middle miocene deposits, zonal definitions and observed assemblages in this study are described in section 9.10.1. the chronostratigraphy of the silicoflagellate zones is established via correlation with the dinocyst stratigraphy of dybkjær & piasecki (2010) and this study (fig. 4). 9.10.1. silicoflagellate zonation lower corbisema triacantha zone (locker & martini 1989) definition. from the lo of naviculopsis quadratum to the fo of mesocena diodon (now known as bachmannocena diodon). age. early to middle miocene (this study, burdigalian to early langhian, cordosphaeridium cantharellus – labyrinthodinium truncatum dinocyst zones, dybkjær & piasecki 2010). observations. in the 2/8–g10a well, the co-occurrence of corbisema triacantha, mesocena apiculata apiculata, octatis speculum ssp. and distephanus crux ssp. (common) without bachmannocena diodon or naviculopsis quadratum indicates the presence of the lower corbisema triacantha zone. assemblages contained octatis speculum hemisphaericum, octatis speculum speculum and distephanus pentagona. siliceous microfossil assemblages were barren with respect to silicoflagellates below 1666.72 m. the lower corbisema triacantha zone is only identified in the 2/8–g10a well (fig. 5, supplementary files s1, s7). a joint upper and lower corbisema triacantha ‘zone’ is assigned to the 2/8–n4, 2/8–8 and 2/11–1 wells (supplementary files s3, s4, s6, s9, s10, s12) as the marker for the top of the lower corbisema triacantha zone, bachmannocena diodon, was not seen in these wells. upper corbisema triacantha zone (locker & martini 1989) definition. from the fo of mesocena diodon (now known as bachmannocena diodon) to the lo of corbisema triacantha. age. middle miocene (this study, late langhian to early serravallian, unipontidinium aquaeductum dinocyst zone, dybkjær & piasecki 2010). observation. in the 2/8–g10a well, the co-occurrence of corbisema triacantha, distephanus crux ssp. (common) and bachmannocena diodon indicates the presence of the upper corbisema triacantha zone. assemblages also contain octatis speculum ssp., mesocena elliptica, mesocena dumitricae, dictyocha aspera, octatis speculum hemisphaericum and octatis speculum quintus. the total range of bachmannocena diodon occurs within the upper corbisema triacantha zone in this study. its lo in the norwegian sea and iceland and rockall plateaux is younger, ranging up into the late miocene to early pliocene (ciesielski et al. 1989; amigo 1999). the upper corbisema triacantha zone is only identified in the 2/8– g10a well (fig. 5, supplementary files s1, s7). a joint upper and lower corbisema triacantha ‘zone’ is assigned to the 2/8–n4, 2/8–8 and 2/11–1 wells (supplementary files s3, s4, s6, s9, s10, s12) as the marker for the base of the upper corbisema triacantha zone, bachmannocena diodon, was not seen in these wells. the base of the overlying paramesocena circulus apiculata zone is based on the lo of corbisema triacantha. this event is seen in several wells but the only silicoflagellate species https://doi.org/10.34194/5k9dv133 http://www.geusbulletin.org/ sheldon et al. 2025: geus bulletin 59. 8381. https://doi.org/10.34194/5k9dv133 21 of 30 geusbulletin.org fig. 8 caption on page 23. n sb 11 n sb 10 n sb 9 1460 1450 2/11 1 1470 1480 1490 1500 1510 1520 1530 1540 1550 1560 1570 1580 1590 1600 1610 1620 1630 1640 1650 1660 1670 1680 1690 1710 1720 1730 1740 1750 1760 1770 1780 1790 1800 1700 g am m a ra y lo glithostratigraphy se qu en ce s c hr on ost ra ti gr ap hy d in oc ys t z on e m ic ro fo ss ils (k in g 20 16 ) sbf1 mfs e sbe mfs d2 sb d2 mfs d1 sbd1 mrs d1 n or dl an d g ro up lo w er m io ce ne bu rd ig al ia n u pp er a qu it an ia n – lo w er m os t b ur di ga lia n lo w er a qu it an ia n la ng hi an se rr av al lia n to rt on ia n m id dl e m io ce ne u pp er m io ce ne ri be g ro up d an y fo rm at io n n or a fo rm at io n m åd e g ro up h od de f or m at io n o �s ho re d en m ar k o �s ho re n or w ay h or da la nd g ro up la rk f or m at io n h om ot . s pp . g. v. a. u. a. a. ca lig od in iu m am ic ul um t. p. – l. s.h. u pp er s um at ra di ni um ha m ul at um c. c./ e. i. c ou st ea ud in iu m a ub ry ae la by rin th od in iu m tr un ca tu m u ni po nt id in iu m aq ua ed uc tu m d ia to m z on e m ic ro fo ss ils (k in g 19 89 ) u nd i� er en ti at ed u nd i� er en ti at ed d en tic ul op sis h ya lin a rh iz os ol en ia b ul bo sa th al as sio sir a fr ag a n sb 12 c n s3 6c -3 8 n s3 839 n sb 12 ab n s3 6a n s3 5b -c n s3 5a n s3 4 2/11 12s fig. 8a 1480 1490 1500 1510 1520 1530 1540 1550 1560 1570 1580 1590 1600 1610 1620 1630 1640 1650 1660 1670 1680 g am m a ra y lo g se qu en ce s d ep th (m ) d in oc ys t z on e m ic ro fo ss ils (k in g 20 16 ) sbf mfs e mrs e sbe mfs d2 sbd2 mrs d1 mfs d1 s. h. g. v. a. a. c. c. c ou st ea ud in iu m a ub ry ae la by rin th od in iu m tr un ca tu m u ni po nt id in iu m a qu ae du ct um d ia to m z on e m ic ro fo ss ils (k in g 19 89 ) u nd iff er en ti at ed u nd i� n s3 6c -3 8 n s3 839 n s3 6a -b n s3 5a n s3 5b -c n sb 12 c n sb 12 ab n sb 10 n sb 11 d en tic ul op sis h ya lin a e. i. t. fr ag a rh iz os ol en ia b ul bo sa c or es 1750 1740 2/8 v6 1760 1770 1780 1790 1800 1810 1820 1830 1840 1850 1860 1870 1880 1890 1900 1910 1920 1930 1940 1950 1960 1970 1980 2000 2010 2020 2030 2040 2050 2060 1990 g am m a ra y lo g se qu en ce s d in oc ys t z on e m ic ro fo ss ils (k in g 20 16 ) sb f1 mfs e sbe mfs d2 sbd2 mfs d1 mrs d1 s. h am ul at um c . ca nt ha re llu s e. i. c ou st ea ud in iu m a ub ry ae la by rin th od in iu m tr un ca tu m u ni po nt id in iu m aq ua ed uc tu m d ia to m z on e m ic ro fo ss ils (k in g 19 89 ) rh iz os ol en ia b ul bo sa th al as sio sir a fr ag a u nd i� . d . h ya lin a n s3 5a n s3 6a -b n s3 5b -c n sb 12 ab n s3 6c -3 8 n sb 10 n sb 11 u. r. n or w eg ic a c . v ig ila ns n . m al ei nt er pr et ar ia r. n or w eg ic a c . v ig ila ns n . m al ei nt er pr et ar ia d ep th (m ) d ep th (m ) gr--> gr--> gr--> a, f c a, b, a a b d g, k i, j, l q a, b c a d e f d g, k h, r j, i, l d, g i, m, t k, m, j, n, q l, k o v n, l m r, w p, q, p r, s t, o b d e, g h k l m o, p q, x r u, x s, y, z aa, w, y e, f d, f, g, h, j, r i, k l, m, t, u, v n, o, p r s t d e c f e m, n o, p, q r s, t, u, v g, h g, d h i l m n q, x r, s, t u, v w, y, w, y z b, j, i, j, k d, g, k, l k m l,n o p, q r, o, p s t, x https://doi.org/10.34194/5k9dv133 http://www.geusbulletin.org/ sheldon et al. 2025: geus bulletin 59. 8381. https://doi.org/10.34194/5k9dv133 22 of 30 geusbulletin.org 2010 2000 2/8 n4 2020 2030 2040 2050 2060 2070 2080 2090 2100 2110 2120 2130 2140 2150 2160 2170 2180 2190 2200 2210 2220 2230 2240 2260 2270 2280 2290 2250 g am m a ra y lo g se qu en ce s d in oc ys t z on e m ic ro fo ss ils (k in g 20 16 ) sbf gr--> mfs e sbe mfs d2 sbd 2 mfs d1 mrs d1 u. s.h. e. i. c . c an t. c ou st ea ud in iu m au br ya e la by rin th od in iu m tr un ca tu m d ia to m z on e m ic ro fo ss ils (k in g 19 89 ) u nd i� . u nd i� er en ti at ed n s3 5a n s3 6c -3 8 n s3 5b -c n s3 6a n sb 11 n sb 12 ab n sb 10 1460 1450 2/8 g10a 1470 1480 1490 1500 1510 1520 1530 1540 1550 1560 1570 1580 1590 1600 1610 1620 1630 1640 1650 1660 1670 1680 1690 1700 1710 1730 1740 1720 g am m a ra y lo g c or es m ic ro fo ss ils (k in g 20 16 ) sbe mfs d2 mrs e sbd2 mfs d1 mrs d1 sbd1 *** a. a. ? ? l. s. h – c. a u pp er s . h am ul at um c . c an th . c ou st ea ud in iu m a ub ry ae la by rin th od in iu m tr un ca tu m u ni po nt id in iu m a qu ae du ct um d ia to m z on e r. n or w eg ic a c . v ig ila ns n itz sc hi a m al ei nt er pr et ar ia th al as sio sir a fr ag a u nd i� er en ti at ed t. fr ag a rh iz os ol en ia b ul bo sa d en tic ul op sis h ya lin a m ic ro fo ss ils (k in g 19 89 ) u nd i� n sb 12 a n sb 11 n sb 10 u nd i� n s3 6c n s3 6a -b n s3 5b n s3 5a fig. 8b sbd 1 se qu en ce s d in oc ys t z on e e. i. ?c. g. nsb9 ns34 ?h. s. 1380 1370 2/8 8 1390 1400 1410 1420 1430 1440 1450 1460 1470 1480 1490 1500 1510 1520 1530 1540 1550 1560 1570 1580 1590 1600 1610 1630 1640 1650 1660 1670 1680 1690 1700 1710 1720 1620 g am m a ra y lo g se qu en ce s d in oc ys t z on e m ic ro fo ss ils (k in g 20 16 ) sbf mfs e sbe mfs d2 sbd 2 mfs d1 mrs d1 lo w er s . h am ul at um u pp er s . h am ul at um e. i./ c. c. c ou st ea ud in iu m a ub ry ae la by rin th od in iu m tr un ca tu m u ni po nt id in iu m a qu ae du ct um g. v. a. a. d ia to m z on e m ic ro fo ss ils (k in g 19 89 ) th al as sio sir a fr ag a rh iz os ol en ia b ul bo sa d en tic ul op sis h ya lin a u nd i� u nd i� er en ti at ed n s3 5a n s3 6c -3 8 n s3 738 u nd i� . u nd i� . n sb 12 b n s3 5b ns35c n s3 4 n sb 9 n sb 12 ab n sb 10 sbd 1 u . aq ua ac ho m os ph ae ra an da lo us ie ns e rh iz os ol en ia b ul bo sa d en tic ul op sis h ya lin a ns37-38nsb12b ns36a-bnsb11 r. n or w eg ic a c . v ig ila ns n . m al ei nt er pr et ar ia d ep th (m ) d ep th (m ) d ep th (m ) gr--> gr--> a, b, c b b a d g l n, o, v a b c, d f a e, e d, g f, j g, k, j, k, r a, b, i, m, o, q, r c d e f g d k h, i, j, g k, j, l, n, u l d, g, h, j, k, p, t, v l, m, n o p, q s, m r, n t, p, r, u, x o, q, s s, w j, k, l, p, t m m n o p r, u, y s w, z n, l o p, q w r s t, q a a, b b c d e h, d i, e, i f, g h k j m, n, o, p q, r, s, t, v w m n o r, v, x, y d, f, c e f, g, h, j, k, l i j, k m, l n o p, q r s, p t, u p, q, p fig. 8 (continued) caption on next page. https://doi.org/10.34194/5k9dv133 http://www.geusbulletin.org/ sheldon et al. 2025: geus bulletin 59. 8381. https://doi.org/10.34194/5k9dv133 23 of 30 geusbulletin.org present above the lo of corbisema triacantha is rare distephanus crux, which cannot be used to confirm the presence of the paramesocena circulus apiculata zone. 9.10.2. silicoflagellate events locker & martini (1989) presented stratigraphically useful early to middle miocene silicoflagellate events from the norwegian sea. some of those events were recognised in this study (figs. 5, 6, supplementary file s1–12) and may be useful for correlation within the valhall–hod area. these are as follows: • the fo of distephanus crux is found in the burdigalian cordosphaeridium cantharellus/sumatradinium hamulatum dinocyst zones in the 2/8–g10a, 2/8–8, 2/8–v6 and 2/11–1 wells. in the 2/11–12s and 2/8–n4 wells, the event occurs in the early langhian labyrinthodinium truncatum zone. the apparent diachroneity of the event in the valhall–hod area may be due to the scarcity of silicoflagellates. • the fo of corbisema triacantha is noted in the lower part of the cousteadinium aubryae dinocyst zone in the 2/8–g10a well and towards the base of the labyrinthodinium truncatum dinocyst zone in the 2/8–n4, 2/8–8 and 2/11–1 wells (supplementary files s3, s4, s6, s9, s10, s12). the apparent diachroneity of this event in the valhall–hod region may be due to the scarcity of silicoflagellates. • the fo of bachmannocena diodon is found locally towards the base of the early langhian unipontidinium aquaeductum dinocyst zone in the 2/8–g10a well (fig. 5, supplementary files s1, s7). it is only noted in the 2/8–g10a cored section but is an important event as it marks the base of the upper corbisema triacantha silicoflagellate zone in the norwegian sea (locker & martini 1989). • the lo of corbisema triacantha is an important event as it marks the top of the upper corbisema triacantha silicoflagellate zone in the norwegian sea (locker & martini 1989). the lo of corbisema triacantha is found below the lo of distephanus crux in the early langhian unipontidinium aquaeductum dinocyst zone in the 2/8–g10a and 2/8–8 wells (fig. 5, supplementary files s1, s4, s7, s10), and lower, in the late langhian labyrinthodinium truncatum dinocyst zone in the 2/8–n4 and 2/11–1 wells (supplementary files s3, s6, s9, s12). the apparent diachroneity of this event in the valhall–hod region may be due to the scarcity of silicoflagellates. • the lo of distephanus crux is diachronous in the valhall–hod area. it is found within the unipontidinium aquaeductum dinocyst zone in the 2/8–g10a, 2/11–12s, 2/8–8, 2/8–v6 and 2/11–1 wells and slightly lower in the late langhian labyrinthodinium truncatum dinocyst zone in the 2/8–n4 well. 10. correlation and stratigraphy the biostratigraphic framework presented here is the result of the analysis of two cored sections (wells 2/11-12s and 2/8-g10a) and four non-cored sections fig. 8 (continued) multidisciplinary biostratigraphic correlation diagram of the six studied wells, from south to north, from the hod field to the valhall field using dinocyst, microfossil and diatom events. events are correlated using coloured correlation lines and are denoted using a letter as follows: dinocysts (purple lines): a: fo sumatradinium hamulatum, b: lo dinopterygium cladoides, c: lo thalassiophora rota, d: fo exochosphaeridium insigne, e: lo cordosphaeridium cantharellus, f: fo cousteaudinium aubryae, g: lo exochosphaeridium insigne, h: fo cerebrocysta poulsenii, i: fo palaeocystodinium miocaenicum, j: fo labyrinthodinium truncatum, k: fo invertocysta tabulata, l: lo cousteaudinium aubryae, m: fo unipontidinium aquaeductum, n: fo habibacysta tectata, o: fo habibacysta tectata (common), p: fo operculodinium erikanium, q: lo cleistosphaeridium placacanthum, r: lo unipontidinium aquaeductum, s: lo sumatradinium hamulatum, t: fo impagidinium sp a wrenn & kokinos 1996 (common), u: fo invertocysta lacrymosa, v: fo cannosphaeropsis passio, w: lo cannosphaeropsis passio, x: lo palaeocystodinium miocaenicum, y: fo achomosphaera andalousiensis, z: fo grammocysta verricula, aa: fo amiculosphaera umbracula. microfossils (blue lines): a: lo plectofrondicularia seminuda, b: lo aulacodiscus allorgei, c: fo globorotalia praescitula, d: fo uvigerina tenuipustulata, e: lo globigerinoides obliquus, f: fo globorotalia zealandica, g: lo trifarina bradyi, h: lo alabamina wolterstorffi, i: lo ceratobulimina contraria, j: lo globorotalia zealandica, k: lo globorotalia praescitula, l: lo uvigerina tenuipustulata, m: lo paragloborotalia nana, n: lo ciperoella ciperoensis, o: lo trilobatus trilobus, p: lo globoturborotalita eolabiacrassata, q: lo elphidium inflatum, r: lo lenticulina peregrina, s: lo cancris subconicus, t: lo loxostomum sinuosum, u: lo sphaeroidinellopsis disjuncta, v: lo asterigerina guerichi staeschei, w: lo bolboforma platyreticulata, x: fo uvigerina kingi, y: lo uvigerina kingi. diatoms (green lines): a: fo high abundance and diversity diatom assemblages, b: fo opephora gemmata, c: fo stictodiscus aff. kittonianus, d: fo thalassionema nitzschiodes (common to abundant), e: fo raphoneis margaritalambata, f: lo raphoneis margaritalambata, g: fo stephanopyxis horridus (common to abundant), h: fo raphoneis robustata, i: fo rhizosolenia miocenca, j: fo cymatosira biharensis, k: fo synedra spp. (influx), l: fo stephanopyxis grunowii, m: lo stephanopyxis horridus (common to abundant), n: fo actinocyclus ingens, o: fo denticulopsis hyalina, p: fo denticulopsis spp. (common to abundant), q: fo denticulopsis hustedtii, r: lo thalassionema nitzschiodes (common to abundant), s: lo diploneis smithii, t: lo high abundance and diversity diatom assemblages (including a. ingens, denticulopsis spp., o. gemmata, s. grunowii, s. aff. kittonianus, synedra spp.). sequence boundary correlations are marked in brown. the datum line for the correlation is sequence boundary sbe. lithostratigraphy for the norwegian sector from eidvin et al. (2022), onshore denmark from rasmussen et al. (2010) and offshore denmark from rasmussen et al. (in press). sequences from rasmussen (2004a; 2017) and dybkjær et al. (2021). gamma (gr) log c/o aker bp. cored intervals are indicated. chronostratigraphy of raffi et al. (2020), dinocyst zonation of dybkjær & piasecki (2010), microfossil zonations of king (1989, 2016) and diatom zonation of schrader & fenner (1976). dinocyst zone abbreviations: a. a.: achomosphaera andalousiensis, a. u.: amiculosphaera umbracula, c. a.: caligodinium amiculum, c. c.: cordosphaeridium cantharellus, c. g.: chiropteridium galea, e. i.: exochosphaeridium insigne, g. v.: gramocysta verricula, h. s.: homotryblium spp., l. s. h.: lower sumatradinium hamulatum, s. h.: sumatradinium hamulatum, t. p.: thalassiphora pelagica, u. s. h.: upper sumatradinium hamulatum, ***: hystrichosphaeropsis obscura, u.: undifferentiated. grey shading denotes no information. a full-sized version of this figure is available as supplementary file s13. https://doi.org/10.34194/5k9dv133 http://www.geusbulletin.org/ sheldon et al. 2025: geus bulletin 59. 8381. https://doi.org/10.34194/5k9dv133 24 of 30 geusbulletin.org (wells 2/8-n4, 2/8-v6, 2/8-8 and 2/11-1). the resulting framework with new and established dinocyst, microfossil, calcareous nannofossil, diatom (small fraction) and silicoflagellate events is presented in figs. 5, 6, 8 and supplementary files s1-s6. the new and established fossil events are shown in supplementary files s7–s12, denoted by coloured correlation lines (dinocysts-purple, microfossils-blue, diatoms-green, sequence boundaries-brown). these correlations are the basis for the conceptual valhall–hod correlation diagram (fig. 8 and supplementary file s13) showing a south to north correlation using dinocyst, microfossil and diatom (small fraction) events and sequence boundaries. the successful application of the biostratigraphic correlation of new and established events in the valhall–hod area implies that they could also be valuable in a more regional context. the studied succession comprises the aquitanian– serravallian (early miocene–late middle miocene) interval of the valhall–hod area. our biostratigraphic study has enabled the dating of the diatomite-rich succession in the valhall–hod area and correlation of the succession with the norwegian lithostratigraphy and the more detailed, newly established lithostratigraphy for the danish sector (rasmussen et al., in press). the study also forms the framework for correlating sequence stratigraphic surfaces and units from the central north sea basin to onshore denmark (rasmussen 2004b, 2017; dybkjær et al. 2021). the thickest diatomite-rich interval within the miocene succession (c. 100 m thick in the 2/8–g10a core on the crest of valhall structure, fig. 5, supplementary file s1; c. 80 m thick in the 2/11–12s core on the crest of hod structure, fig. 6, supplementary file s2) is of langhian age and correlates with the nora formation. the nora formation was defined in the danish sector of the north sea (rasmussen et al., in press) and correlates with the upper part of the odderup formation and lower part of the hodde formation defined for onshore denmark (fig. 3). the underlying minor diatomite-rich intervals are of burdigalian age and correlate with the danish dany formation (rasmussen et al., in press). the bastrup formation and the lower and upper odderup formation, respectively, are defined onshore denmark (see fig. 3). the nora and odderup formations correlate with the uppermost part of the lark formation defined in the norwegian sector of the north sea, while the dany formation and the time-equivalent lithostratigraphic units defined for onshore denmark correlate with the upper part of the lark formation (rasmussen et al. in press; fig. 3). the combination of biozones and additional events provides a stratigraphic subdivision of the diatomiterich lower to middle miocene succession in the order of 5–15 m in the fully cored 2/11–12s well, somewhat less in the 2/8–g10a well (due to the core gap) and in the non-cored wells. this extremely high-resolution subdivision forms a solid basis for improved reservoir evaluation. no major hiatuses are recognised, though there appears to be a potential repeated section in the uppermost part of the core in the 2/8–g10a well. 11. discussion 11.1. improvements to regional and local biostratigraphy the multidisciplinary biostratigraphy, with numerous microfossil events presented in this study, is useful for reservoir subdivision and reservoir characterisation in the valhall–hod area and for local and regional correlation. dinocysts were found in all samples in this study. the zonation of dybkjær & piasecki (2010) was used successfully. all the events found in their study and 15 new dinocyst events are recognised. fewer events are found in the interval where the highest content of silica occurs in the nora formation equivalent (l. truncatum zone). the dinocyst study of the valhall–hod area has resulted in the improvement of the zonation of dybkjær & piasecki (2010). the microfossil zonations of king (1989, 2016) are applied to the studied sections. twenty-five microfossil events are recognised in this study (fig. 8 and supplementary file s13), and most are concentrated in the upper part of zone nsb10 (ns35b-c) and in nsb11 (ns36a-b). nannofossils are not commonly used for biostratigraphy in the north sea miocene interval due to the lack of calcareous sediments at this stratigraphic level and the successful application of dinocysts and microfossils. however, the global zonation of martini (1971) is successfully applied to the low-resolution nannofossil study and supports the biostratigraphy based on other microfossil groups. seventeen nannofossil events are recognised in this study that have correlation potential (fig. 7). most of these nannofossil events are documented from the north sea area for the first time in this study. they have been previously described from the mediterranean area and the low latitudes (young 1998) and the gulf of mexico (de kaenel et al. 2017; browning et al. 2017; boesiger et al. 2017) and are potentially important for correlation beyond the north sea basin. the detailed study of siliceous microfossils from the fine fraction of the cored wells 2/8–g10a (valhall field) and 2/11–12s (hod field) has resulted in the recognition of a series of diatom and silicoflagellate events that correlate between the two cored wells situated on two neighbouring structures. many of these siliceous microfossil events https://doi.org/10.34194/5k9dv133 http://www.geusbulletin.org/ sheldon et al. 2025: geus bulletin 59. 8381. https://doi.org/10.34194/5k9dv133 25 of 30 geusbulletin.org are also found in nearby non-cored wells: 2/8–n4, 2/8–v6 and 2/8–8 (valhall) and 2/11–1 (in the saddle between the valhall and hod fields). this highlights the diatom correlation potential in a local context and that they could form the basis of a north sea siliceous microfossil zonation at a later stage. while not all the diatom species that mark the tops and bases of the zones of schrader & fenner (1976) are found in this study, the diatom assemblages from the valhall–hod area compare well with those from the norwegian sea. the diatom biostratigraphy and 20 diatom events described in this study are useful for the subdivision of the diatomite or silica-rich nora formation equivalent and dany formation equivalent (uppermost part of the lark formation) and potentially a valuable tool for hydrocarbon reservoir characterisation. the diatom biostratigraphy can potentially be used to correlate regionally, from the north sea farther north into the norwegian sea. this is the first time, potentially, that siliceous microfossils have been used to correlate within the miocene succession of the north sea basin and regionally. silicoflagellate identification allowed the recognition of two zones of locker & martini (1989) and five stratigraphically significant events are recognised in this study. delicate silicoflagellate skeletons in our samples were often fragmented, perhaps due to harsh conditions the samples were subjected to during preparation. we are presently testing a less destructive method, which will probably benefit future silicoflagellate studies. diatoms and silicoflagellate biostratigraphic events have been discussed in detail in this study. other siliceous components (e.g. sponge spicules, radiolaria and ebridians) are also prevalent in parts of the lower to middle miocene ‘diatomite’ successions. further detailed studies concentrating on the individual siliceous elements, their relative abundances and palaeoecology are currently underway, which will hopefully shed light on the conditions and driving forces necessary for mass siliceous microfossil production, which, in turn, will potentially aid prediction of siliceous reservoir properties and reservoir correlation across the valhall–hod area. for the early to middle miocene diatomite-rich interval of the valhall–hod area, a combination of dinocyst, microfossil and diatom biostratigraphy appears to be particularly promising for detailed subdivision of the siliceous reservoir interval (fig. 8 and supplementary file s13). 11.2. sequence stratigraphic and lithostratigraphic correlation this robust biostratigraphic framework produced in this study allows correlation of the sequence stratigraphy defined onshore denmark to the central parts of the north sea basin as presented in rasmussen (2004, 2017) and dybkjær et al. (2021) and also to the new danish offshore lithostratigraphy for the neogene (rasmussen et al., in press). 11.3. age model a solid age model for the miocene does not exist for the north sea area. however, steps have been taken to attain this goal. eidvin et  al. (2014b) used sr isotopes to analyse the whole of the miocene section from samples from jylland, onshore denmark. results from that study showed that sr dating of samples from the lower miocene and lower middle miocene succession supports the datings of the dinocyst zonation of dybkjær & piasecki (2010). however, in the upper middle miocene and upper miocene parts of the section, the sr ages are too old compared with ages based on dinocysts and bolboforma. the authors of these publications agree that there are discrepancies with the sr dating. the lack of a robust age model for the miocene of the north sea area is also a major reason for the use of a selection of vintage, but strongly reliable, biostratigraphic zonations in this study. we considered it important to correlate fossil zones and events, rather than try to tackle the chronostratigraphic problem. the five biostratigraphic zonation schemes in fig. 4, one for each fossil discipline, are correlated with the chronostratigraphic time scale that was relevant at the time of publication of the zonation. our study has provided a direct, reliable correlation for the lower and middle miocene succession of the valhall–hod area, based on five fossil groups using a series of events recognised in a selection of closely spaced wells, mostly on the same samples. we acknowledge the value of the neogene timescale of raffi et al. (2020), which contains dinocyst, foraminifera, calcareous nannofossil and diatom data and events from a multitude of global localities. however, our study is centred in the north sea area, which was a semi-enclosed basin during the miocene period. miocene north sea microfossil assemblages and events share more affinities to northern midto high-latitude and boreal assemblages than those from the low latitude and tropical locations cited in raffi et  al. (2020). we conclude that a reliable and robust palaeomagnetic time frame for the north sea area would be valuable for future stratigraphic studies. 11.4. palaeoclimate the studied interval in this project includes the mco (c.17–13 ma; e.g. larsson et  al. 2011; herbert et  al. 2020; sliwinska et al. 2024) and the mmct (c. 14.7–13–8 https://doi.org/10.34194/5k9dv133 http://www.geusbulletin.org/ sheldon et al. 2025: geus bulletin 59. 8381. https://doi.org/10.34194/5k9dv133 26 of 30 geusbulletin.org ma; flower & kennett 1994). the high-resolution biostratigraphic framework thus forms a good basis for future studies of the climatic and palaeoenvironmental changes during these important time intervals. 12. conclusions a new multidisciplinary biostratigraphic framework is established using core samples from the lower-middle miocene succession of the 2/8–g10a and 2/11–12s wells from the valhall and hod fields, respectively. the core study provides a correlation using five biostratigraphic disciplines: dinocysts, foraminifera, calcareous nannofossils, diatoms and silicoflagellates. the correlation is based mostly on the same samples, providing a unique and exceptionally detailed, robust framework. this framework was successfully tested on the equivalent chronostratigraphic level of the 2/8–n4, 2/8–v6, 2/8–8 (valhall field) and 2/11–1 (in the saddle between the valhall and hod fields) wells based on ditch cutting samples. the studied interval spans the chiropteridium galea zone to the hystrichosphaeropsis obscura dinocyst zone (dybkjær & piasecki 2010), the nsb9 (ns34) to nsb12c (ns38–39) microfossil zones (king 1989, 2016), nannofossil zones nn3–nn6 (martini 1971), the rhizosolenia norwegica–coscinodiscus vigilans–nitzschia maleinterpretaria joint prz to the denticulopsis hyalina prz (diatoms, schrader & fenner 1976) and the lower and upper corbisema triacantha silicoflagellates zones (locker & martini 1989). twenty-seven dinocyst events, 25 microfossil events, 17 nannofossil events, 20 diatom events and 5 silicoflagellate events are recognised for the studied interval. the new dinocyst, microfossil, nannofossil and diatom events (bases, tops and occurrences) from this study are successfully used to correlate on and between the valhall field and hod field and can be used to supplement published zonation schemes and events. the successful application of the new framework in the valhall–hod area implies that it could also be valuable in a more regional context. a suite of diatom events is recognised for the lower and middle miocene succession of the valhall–hod area and potentially forms the basis for a diatom zonation for the north sea basin. to the best of our knowledge, this diatom study (fine fraction) is the first of its kind covering the early and middle miocene interval of the north sea and is potentially useful for regional correlation and local reservoir characterisation. our low-resolution calcareous nannofossil study of the valhall–hod part of the north sea basin recognises events described from low-latitude and tropical locations. the siliceous or diatomite-rich reservoir interval has been dated, and a detailed (5–15 m interval) biostratigraphic subdivision is provided. the detailed biostratigraphy has enabled the correlation of the sequence stratigraphic surfaces and lithostratigraphic units defined in the danish sector to the southern norwegian sector. acknowledgements aker bp and pandion energy are acknowledged for their initiation and funding of the study, as well as for their valuable contributions through discussions and the exchange of ideas throughout the project’s duration. special thanks go to geus laboratory technicians annette ryge and charlotte olsen for processing a huge number of samples. jacob lind bendtsen is thanked for his unwavering patience when preparing the figures. reviewers haydon bailey and erik anthonissen are thanked for their comments, which helped to develop the manuscript for publication. additional information funding statement aker bp and pandion energy funded this study. author contributions es and kd: conceptualisation, writing – original draft, writing – review and editing. es: biostratigraphy (microfossils, calcareous nannofossils, diatoms, silicoflagellates). kd: biostratigraphy (dinocysts). esr: lithostratigraphy, sequence stratigraphy. mo: initial diatom supervision and biostratigraphy (diatoms). competing interests the authors declare no competing interests. additional files thirteen supplementary files are available at https://doi.org/10.22008/ fk2/vlo4ln: supplementary files s1–s6.pdf contain fig. s1 biostratigraphic summary diagram for the 2/8–g10a well, fig. s2 biostratigraphic summary diagram for the 2/11–12s well, fig. s3 biostratigraphic summary diagram for the 2/8–n4 well, fig. s4 biostratigraphic summary diagram for the 2/8–8 well, fig. s5. biostratigraphic summary diagram for the 2/8–v6 well and fig. s6 biostratigraphic summary diagram for the 2/11–1 well. supplementary files s7–s12.pdf include correlation lines fig. s7 biostratigraphic summary diagram for the 2/8–g10a well, fig. s8 biostratigraphic summary diagram for the 2/11–12s well, fig. s9 biostratigraphic summary diagram for the 2/8–n4 well, fig. s10 biostratigraphic summary diagram for the 2/8–8 well, fig. s11 biostratigraphic summary diagram for the 2/8–v6 well and fig. s12 biostratigraphic summary diagram for the 2/11–1 well. supplementary file s13.pdf contains fig. s13 multidisciplinary biostratigraphic correlation diagram (full size) of the six studied wells using dinocyst, microfossil and diatom events. references amigo, a.e. 1999: miocene silicoflagellate stratigraphy: iceland and rockall plateaus. in: raymo, m.e. et al. 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(ed.): 2nd. ed. shell international petroleum, 1–239, geological society publication. london. https://doi.org/10.34194/5k9dv133 http://www.geusbulletin.org/ https://doi.org/10.2973/odp.proc.sr.189.107.2004 https://doi.org/10.4095/103330 https://doi.org/10.58998/jnr2294 https://doi.org/10.1007/978-94-011-4902-0_8 https://doi.org/10.1007/978-94-011-4902-0_8 https://www.mikrotax.org/nannotax3 https://www.mikrotax.org/nannotax3 http://mikrotax.org https://www.mikrotax.org/pforams/index.php?dir=pf_cenozoic https://doi.org/10.1126/science.1059412 a multidisciplinary biostratigraphic framework for the lower to middle miocene of the norwegian nort 1. introduction 2. geological setting and palaeoclimate 3. lithostratigraphy 4. sequence stratigraphic framework 5. absolute dating 6. previous studies 7. zonation schemesin this study 8. materials and methods 8.1. preparation methods 8.1.1. palynology 8.1.2. microfossils (large fraction) 8.1.3. calcareous nannofossils 8.1.4. siliceous microfossils (small fraction) 9. results 9.1. palynology 9.2. dinocyst events 9.3. microfossils 9.4. microfossil events 9.5. calcareous nannofossils 9.6. calcareous nannofossil events 9.7. diatoms 9.8. diatom zonation 9.9. diatom events and observations from this study 9.10. silicoflagellates 9.10.1. silicoflagellate zonation 9.10.2. silicoflagellate events 10. correlation and stratigraphy 11. discussion 11.1. improvements to regional and local biostratigraphy 11.2. sequence stratigraphic and lithostratigraphic correlation 11.3. age model 11.4. palaeoclimate 12. conclusions acknowledgements additional information funding statement author contributions competing interests additional files references figures fig. 1 palaeogeography of the north sea area in the early miocene. the red dot indicates the locati fig. 2 depth map to the top miocene of the valhall and hod structures and locations of the six studi fig. 3 the new lithostratigraphic subdivision for the neogene succession in the danish sector of the fig. 4 (continued) summary diagram for the 2/8-g10a well with biostratigraphic zones, previously pub fig. 5 summary diagram for the 2/11-12s well with biostratigraphic zones, previously published event fig. 8 multidisciplinary biostratigraphic correlation diagram of the six studied wells, from south larsen et al. 2024: geus bulletin 57. 8361. https://doi.org/10.34194/geusb.v57.8361 1 of 27 research article a hidden succession revealed: cretaceous and paleocene sediments and a native-iron-bearing lava flow in cores near qullissat, disko, west greenland lotte melchior larsen*1 , gunver krarup pedersen1 , henrik nøhr-hansen1 , asger ken pedersen1,2, jørgen a. bojesen-koefoed1 , erik vest sørensen1 , sofie lindström1,3 1geological survey of denmark and greenland (geus), copenhagen, denmark, 2natural history museum of denmark, copenhagen, denmark, 3department of geosciences and natural resource management, copenhagen university, copenhagen, denmark abstract over long stretches of the north-east coast of qeqertarsuaq (disko), the sediments in the nuussuaq basin and their relations to the volcanic rocks are concealed beneath numerous landslides. two cores south of qullissat drilled by falconbridge ltd in 1994, targeting a native-iron-bearing igneous body assumed to be a sill, present well-preserved sections through the hidden succession. we have dated the sediments in the cores palynologically. the lower part comprises 115 m of deltaic deposits, including coal seams, of the cretaceous atane formation, qilakitsoq member (late turonian to early coniacian age), which has not been recorded on disko before. the two cores and five short coastal cliff sections are mutually correlatable and correlate further to the coal seams earlier mined at qullissat; the coals are hereby dated for the first time. the cretaceous rocks are overlain by 15 m of marine deposits, mainly mudstones, of the danian eqalulik formation, with a hiatus of c. 24 million years. the igneous body of native-iron-bearing basaltic andesite has a thick, red-oxidised, vesiculated and brecciated top zone and is interpreted as a subaerial lava flow belonging to the asuk member of the vaigat formation. the flow has run perhaps up to 20 km from the eruption site to the sea, where it ponded and attained a thickness of 138 m, the thickest lava flow in the west greenland basalt group. the flow is overlain by 22 m of non-marine sandstones and mudstones of the atanikerluk formation. the core correlation indicates the existence of a fault with c. 90 m vertical displacement between the two drill sites. the structural relations of the various parts of the atane formation along the north-east coast of disko necessitate the assumption of another hidden, prevolcanic fault south of qullissaaqqat. *correspondence: lml@geus.dk received: 29 aug 2023 revised: 04 jan 2024 accepted: 30 jan 2024 published: 28 mar 2024 keywords: nuussuaq basin, atane formation, qilakitsoq member, vaigat formation, asuk member, palynology, landslides abbreviations a.s.l.: above sea level cu: coarsening upwards geus: geological survey of denmark and greenland lo: last occurrence ro: vitrinite reflectance tc: total carbon tmax: temperature of maximum pyrolysate-yield toc: total organic carbon ts: total sulphur utm: universal transverse mercator geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: karen dybkjær (geus, denmark) reviewed by: henrik tirsgaard (totalenergies, denmark) and james b. riding (british geological survey, uk) funding: see page 25 competing interests: see page 25 additional files: none 1 introduction the nuussuaq basin in west greenland contains exposed sedimentary rocks of cretaceous to palaeogene age overlain and intruded by volcanic rocks of palaeogene age. the basin is one of the most intensely investigated regions in greenland, and coal deposits in the sediments and mineralisations of iron, nickel, copper and platinum-group elements in the volcanic rocks have been explored and exploited for more than hundred years. nowhere in the basin have these activities been more intense than in the northern part of qeqertarsuaq (disko), particularly in the qullissat area where a coal mine and an associated town (qullissat) with more than thousand inhabitants were situated (fig. 1). the geological relations in the nuussuaq basin are generally well mapped and understood (e.g. clarke & pedersen 1976; dam et al. 2009; larsen et al. 2016; pedersen et al. 2017, 2018; dam & sønderholm 2021; larsen & larsen 2022). however, the north-east coast of disko, including the area around qullissat, has been difficult to examine geologically because of numerous landslides (e.g. svennevig et al. 2023). here, cretaceous strata with coal seams of the atane formation are well exposed in the coastal cliffs up to c. 50 m altitude, and the paleocene volcanic succession of the vaigat and https://doi.org/10.34194/geusb.v57.8361 https://orcid.org/0000-0002-9344-4166 https://orcid.org/0000-0002-0792-2257 https://orcid.org/0000-0002-9291-8104 https://orcid.org/0000-0001-5647-2769 https://orcid.org/0000-0002-1827-2894 https://orcid.org/0000-0001-8278-1055 mailto:lml@geus.dk https://creativecommons.org/licenses/by/4.0/deed.ast larsen et al. 2024: geus bulletin 57. 8361. https://doi.org/10.34194/geusb.v57.8361 2 of 27 geusbulletin.org maligât formations is excellently exposed in the steep mountain slopes above c. 700 m altitude. however, the intervening interval at c. 50–700 m altitude is dominated by landslides, and the rocks are extensively covered by talus and vegetation (figs 1 and 2). therefore, the lithostratigraphy of the cretaceous and paleocene sedimentary deposits and the relations between the sedimentary and volcanic rocks are poorly known over long stretches of the north-east coast of disko. even the age of the mined coal seams has not been well established. the occurrences of volcanic rocks with native iron on disko are world-famous (pedersen et al. 2017, 2018 and numerous references therein) and have been repeatedly explored by mining companies (see next section). a native-ironand sulphide-bearing, sill-like igneous body occurs in the sediments between qullissat and qullissaaqqat but is poorly exposed and partly displaced by landslides (fig. 1b). this body has attracted mining companies that assumed an analogy to iron-bearing igneous rocks in the noril’sk region in siberia (lightfoot & hawkesworth 1997, their figs 14 and 15), where a crucial mechanism is considered to be metal and sulphide precipitation during magma flowage through subsurface intrusions (sills). in 1993 and 1994, the mining company falconbridge ltd. drilled five boreholes targeting the supposed sill; most were short due to technical difficulties, but the two cores fp94-4-5 and fp94-4-6 together constitute a long section of excellent quality through cretaceous and palaeogene sedimentary strata as well as the complete igneous body (olshefsky & jerome 1994; olshefsky et al. 1995; lode et al. 2021). this paper presents the results of a study of the two cores, which represent the poorly known stratigraphic interval concealed by landslides. the cores have enabled us to place constraints on the stratigraphy, age and structure of the cretaceous and paleocene sedimentary strata in north-eastern disko; they also show that the native-iron-bearing igneous body is not a sill but a subaerial lava flow, which we assign to the volcanic asuk member of the vaigat formation. 70°n 69°30'n 55°w 53°w 70°30'n nuussuaq disko paatuut kingittoq atanikerluk kussinerujuk asuk qullissat ui�aq ataata kuua ikorfat inussuk qullissaaqqat nuugaarsuk pingu vaigatkuugannguaq b ice cap west greenland basalt group nuussuaq group precambrian basement landslips snow, ice maligât fm vaigat fm tertiary cretaceous k t ls volcanics sediments 25 km 55°w 53°w vaigat fm qullissat qullissaaqqat ba fp94-4-5 fp94-4-6 greenland a 53° 70°04 inussuk fig. 1 geological maps of parts of the nuussuaq basin. a: overview map of disko and nuussuaq, with names mentioned in the text and the position of fig. 1b indicated. b: map of the area around qullissat and qullissaaqqat, with boreholes fp94-4-5 and fp94-4-6 indicated. two exposures of the iron-bearing igneous body studied here are labelled ‘fe’. other labels are part of the geological divisions on the map. excerpt from the geological map 1:100 000 qutdligssat (old spelling, rosenkrantz et al. 1976), reproduced to scale, i.e. 1 cm = 1 km. https://doi.org/10.34194/geusb.v57.8361 http://www.geusbulletin.org/ larsen et al. 2024: geus bulletin 57. 8361. https://doi.org/10.34194/geusb.v57.8361 3 of 27 geusbulletin.org 2 history in 1871 and 1872, k.j.v. steenstrup was sent to greenland with instructions to investigate the occurrence of native iron at uiffaq in south-western disko and at the same time to collect information on the occurrences of coal seams in the region. the second task brought him to the north-east coast of disko, where several coal seams in the atane formation are intermittently exposed in the coastal cliffs between qullissat and qullissaaqqat (gamle qutdligssat) about 3 km south-east of qullissat (figs 1 and 2). the coal had been quarried since the 1770s near qullissaaqqat (‘ritenbenk’s coal quarry’) for local use in the settlements in the qeqertarsuup tunua (disko bugt) region (schiener 1976; shekhar et al. 1982), and the danish authorities wanted to increase the utilisation of the coal in west greenland. steenstrup (1874) published a description and a drawn section of the coal-bearing sediments in the coastal cliffs between the coal quarry and the alluvial fan where the qullissat town later was built. the coal was mined at qullissat between 1924 and 1972, when the mine was closed and the town abandoned (schiener 1976). steenstrup also discovered a native-iron-bearing lava flow at asuk 19 km north-west of qullissat, the first iron occurrence found outside the uiffaq locality. at uiffaq, the origin of the iron, whether meteoritic or telluric, was disputed, and the asuk locality has become famous because it proved beyond doubt that the iron is of telluric origin (steenstrup 1875, 1877, 1882). steenstrup also found the igneous body, which is described in the present work, and sampled it in a landslide at nunngarut (figs 2 and 3); he noted that it was graphite-bearing and only later found that it also carries native iron and closely resembles the iron-bearing rock at asuk (steenstrup 1900, p. 268). more recently, exploration for nickel, copper and precious metals in the volcanic rocks was carried out by the mining companies kryolitselskabet øresund a/s in the 1950–1960s, greenex ltd. in the 1980s and falconbridge ltd. in the 1990s. 3 geological setting the nuussuaq basin is a rift basin that extends for c. 400 km along the coast of west greenland (chalmers et al. 1999; chalmers & pulvertaft 2001; hopper et al. fig. 2 oblique aerial photograph of the coastal area of disko between the closed coal mine (at right) and qullissaaqqat (immediately left of the picture). ritenbenk’s old coal quarry is clearly seen from the air; seen on the ground it is inconspicuous. the positions of boreholes fp94-4-5 and fp94-4-6 and coastal sections q-1, q-2, q3, m-1 and m-2 are indicated. exposures of the nunngarut igneous body (fe) in situ are annotated with white lines to the right and left; a third exposure is seen in a landslide at the coast at nunngarut. the height of the coastal cliffs is around 50 m, and the height of the plateau top of the inussuk mountain is around 900 m. note the tumbled and landslipped surface and lack of exposures between inussuk and the coastal cliffs. compare fig. 3. photo: kristian svennevig 2019. fp94-4-5 ritenbenk’s coal quarry coal mine (closed)m-1q-1 q-2 q-3 m-2nunngarut inussuk fp94-4-6 fe fe fe https://doi.org/10.34194/geusb.v57.8361 http://www.geusbulletin.org/ larsen et al. 2024: geus bulletin 57. 8361. https://doi.org/10.34194/geusb.v57.8361 4 of 27 geusbulletin.org 2016; dam & sønderholm 2021). in the disko–nuussuaq area (fig. 1a), the sedimentary strata are of albian to paleocene age and constitute the nuussuaq group (fig. 4; dam et al. 2009; pedersen & nøhr-hansen 2014; nøhr-hansen et al. 2016; pedersen et al. 2023). the structural development includes an early rift phase, a subsidence phase, a late rift phase and a drift phase. these four phases comprise eight tectonostratigraphic sequences (tss1–tss8) described and interpreted by dam & sønderholm (2021, their fig. 3), see fig. 4. tss1 (albian) comprises fluvial and lacustrine deposits of the kome and slibestensfjeldet formations in northern nuussuaq. during the deposition of tss2 to tss4 (late albian to earliest campanian), a huge delta system built out from the south-east, and sedimentary material was transported towards north-west, west and south-west across a delta plain with coal swamps and shallow lakes (pedersen & pulvertaft 1992). the fluvial system may have had a drainage area that covered large parts of ‘the interior’ of greenland (olsen 1993; dam et al. 2020). the delta front prograded repeatedly towards the north, north-west and west (dam et al. 2009; dam & sønderholm 2021). the fluvial, delta plain, shoreface and delta front deposits constitute the atane formation, which is seen in outcrops in disko and in southern, central and northern nuussuaq. tss4 also includes the lower part of the marine itilli formation, i.e. the kussinerujuk member in northern disko from asuk and westwards (dam et al. 2009). the atane formation is overlain by submarine slope, fan and channel deposits in northern and central nuussuaq. they constitute the aaffarsuaq member of the itilli formation (dam et al. 2009) and belong to tss5 (campanian to early maastrichtian). in central nuussuaq, the base of tss5 is an angular unconformity (dam et al. 2000). this early campanian unconformity is recognised offshore in regional geophysical datasets (gregersen et al. 2013, 2019, 2022). tss6 comprises the kangilia formation (latest maastrichtian and danian age, fig. 4), which consists of a thick conglomerate unit interpreted as a turbidity channel deposit, overlain by deep-water marine turbiditic mudstones (figs 76, 95 in dam et al. 2009). the formation is exposed in northern and western nuussuaq. at the south coast of nuussuaq (ataata kuua), the kangilia formation fills a submarine canyon (dam et al. 2009, figs 14, 15, 91). tss7 (danian) comprises the complex fill of incised valleys of fluvial and submarine origin. thick fluvial and lacustrine successions of the quikavsak formation truncate the atane and kangilia formations in southern fig. 3 photogrammetrically measured and interpreted section along the coast of disko between qullissat (at right) and qullissaaqqat (immediately left of the section). exposures of the iron-bearing nunngarut igneous body are red. the vertical extents of boreholes fp94-4-5 and fp94-4-6 are indicated. compare fig. 2, but note that the perspectives are very different because the section possesses no single point of view. modified from fig. 48 in pedersen et al. (2018) or fig. 161 in pedersen et al. (2017). mm 0 km ls ls ls ls ls ls ls ls ls ls ls ls ls ls ls ls t? t? ls fe fe fe fe fe 1 km 2 km nunngarut coal mine inussuk qullissat 3 km 4 km 5 kmq-3 m-2 m-1q-1 q-2 fe native iron ls landslipped area houses (qullissat) borehole extent vaigat formation, asuk member t? paleocene? cretaceous fluvial sandstone atane formation, sandstone with shale beds and coal layers paleocenequaternary undifferentiated maligât fm subaerial lava flows iron-bearing magnesian andesite foreset-bedded hyaloclastites lower rinks dal mb lava flows with prominent entablatures and thick top breccias, partly or wholly subaqueous (unit 509) boundary between massive lower part and thick top breccia of lava flow skarve�eld unit (unit 511) olivineand feldspar-microphyric basalt ice and perennial snow undifferentiated deposits maligât formation, rinks dal member 200 400 600 800 1000 1200 1400 1600 1800 200 400 600 800 1000 1200 1400 1600 1800 fp94-4-6 fp94-4-5 https://doi.org/10.34194/geusb.v57.8361 http://www.geusbulletin.org/ larsen et al. 2024: geus bulletin 57. 8361. https://doi.org/10.34194/geusb.v57.8361 5 of 27 geusbulletin.org nuussuaq, while thick, marine turbiditic sandstones and conglomerates of the agatdal formation overlie the kangilia formation in central nuussuaq (dam et al. 2009, figs 109, 114). the dramatic changes in water depth from the marine deep-water kangilia formation to the incised fluvial valleys (quikavsak formation) and the return to deep-water conditions (eqalulik formation) at the south coast of nuussuaq indicate a danian uplift presumably caused by arrival of the proto-icelandic mantle plume (dam & nøhr-hansen 2001; dam & sønderholm 2021). tss8 (danian to selandian) comprises both synvolcanic sediments, which drape the prevolcanic deposits, and volcanic rocks. the eqalulik formation in the north and west consists of tuffaceous mudstones deposited in a deep marine embayment bounded to the west by the advancing volcanic front. the atanikerluk formation, in the east and south, consists of siliciclastic sediments deposited from the east in lacustrine and fluvial environments. the lacustrine mudstones include thin layers of volcanic ash (pedersen et al. 1998; dam et al. 2009). the volcanic rocks of tss8 in the disko – nuussuaq area comprise the vaigat and maligât formations (fig.  4). eruption of the vaigat formation magmas started in western nuussuaq and expanded eastwards and southwards to cover successively larger areas with hyaloclastites and overlying associated subaerial lava flows (pedersen et al. 2017). three major volcanic episodes gave rise to the three main members: the anaanaa, naujánguit and ordlingassoq members (fig.  4). the eastward progradation of the hyaloclastites of the anaanaa and naujánguit members into the marine embayment gradually narrowed the connection to the open sea north of nuussuaq. the connection was finally blocked at the end of the second volcanic episode, leaving a brackish and eventually freshwater lake that covered large areas of south-central to south-eastern nuussuaq and north-eastern disko. in this lake, the atanikerluk formation was deposited (pedersen et al. 1996; pedersen et al. 1998). hyaloclastites of the third volcanic episode continued the eastward progradation into the lake and reduced its extent and water depth. the lake expanded westwards during the pause between the vaigat and maligât formations, and when the maligât formation started to erupt from centres in south-western disko, a new lake covered eastern disko and southernmost nuussuaq. this lake was gradually filled with more sediments of the atanikerluk formation (assoq member) and hyaloclastites and subaqueous lava flows of the maligât formation (dam et al. 2009; pedersen et al. 2018, fig. 16). 4 geology of the formations represented in the drill cores 4.1 sedimentary rocks 4.1.1 atane formation the atane formation comprises fluvial and deltaic deposits of late albian to latest santonian–earliest campanian age in nuussuaq and disko (fig. 4). the lower boundary is exposed at the north coast of nuussuaq. in disko, the contact between weathered basement and fluvial deposits is recorded in borehole fp93-3-1 in the kuugannguaq valley west of qullissat (dam et al. 2009, fig. 19). the upper boundary of the atane formation consists of erosional unconformities overlain by the aaffarsuaq member of the itilli formation (central nuussuaq), the quikavsak formation (south coast of nuussuaq), the kussinerujuk member of the itilli formation (north coast of disko) and the eqalulik and atanikerluk formations (northern and eastern disko and south-eastern nuussuaq). the atane formation comprises four members, the distribution of which is shown in fig. 5. the ravn kløft member (late albian age) is known only from the north coast of nuussuaq (fig. 5). it is interpreted as estuarine and is overlain by the deltaic kingittoq member (midtgaard 1996; dam et al. 2009; pedersen & nøhr-hansen 2014; pedersen et al. 2023). fig. 4 stratigraphy and age of the sedimentary and volcanic formations of the nuussuaq basin on disko and nuussuaq. on the age scale to the left, rings around two numbers and change of colours indicate changes in the scale. tss-1 to tss-8 are tectonostratigrahic sequences (dam & sønderholm 2021). ku.: kussinerujuk member. members of the volcanic formations are a: anaanaa member; n: naujánguit member; o: ordlingassoq member; rdm: rinks dal member. albian cenomanian kome formation slibestens�eldet formation ravn kløft member kingittoq member skansen member a ta ne f or m at io n turonian coniacian santonian campanian maastrichtian danian selandian tss8 tss7 tss6 tss5 tss4 lo w er c re ta ce ou s u pp er c re ta ce ou s pa le oc en e ku. qilakitsoq member itilli formation quikavsak fm agatdal fm kangilia formation hiatus eqalulik formation atanikerluk fm vaigat fm rdm n a o maligât formation ma tss3tss3 tss2 tss1110 113 105 100 9595 90 85 80 75 7070 65 64 63 62 61 60 https://doi.org/10.34194/geusb.v57.8361 http://www.geusbulletin.org/ larsen et al. 2024: geus bulletin 57. 8361. https://doi.org/10.34194/geusb.v57.8361 6 of 27 geusbulletin.org the skansen member (late albian to mid-cenomanian age) is exposed in eastern disko (fig. 5). it is characterised by thick fluvial sandstones interbedded with thinner units of mudstone and thin coal beds, interpreted as flood-plain deposits (koppelhus & pedersen 1993). the member is exposed on the north coast of disko from pingu and c. 10 km westwards to nuugaarsuk. from nuugaarsuk and north-westwards to qullissaaqqat, the coastal cliffs are strongly affected by landslides, and the cretaceous deposits form loose, sand-dominated scree slopes, which have not been closely investigated. the sedimentological model predicts that the skansen member is replaced basinwards (north-westwards) by delta deposits of the same age, referred to the kingittoq member (fig. 4). a very long hiatus separates the skansen member from the overlying synvolcanic, fluvial sediments of the paleocene akunneq member of the atanikerluk formation (pedersen et al. 1998; dam et al. 2009, fig. 16). the kingittoq member represents stacked deltaic deposits exposed at the north coast of nuussuaq east of ikorfat, at the south coast of nuussuaq from atanikerluk to kingittoq and at the north coast of disko from east of asuk (‘asuk east’) and westwards to kussinerujuk (fig. 4). between asuk east and kussinerujuk, the erosional unconformity tss3 forms the boundary between the kingittoq member and the overlying kussinerujuk member (itilli formation) deposited in the early part of tss4 (dam et al. 2009; fig. 5). the kingittoq member ranges in age from late albian at asuk, atanikerluk and east of ikorfat to middle turonian at kingittoq (pedersen & nøhr-hansen 2014; pedersen et al. 2023). marine dinoflagellate cysts are scarce in the older part of the member but more frequent in its younger part. a similar pattern is known from the baffin bay (nøhr-hansen et al. 2021). the qilakitsoq member represents stacked deltaic deposits in central nuussuaq and at the south coast from paatuut and westwards (fig. 5). many sections in central nuussuaq show thick delta front deposits, wave-generated sedimentary structures are common and transgressive shoreface sandstones are well developed. the member is of early coniacian to latest santonian age in the borehole ggu 247801 and adjacent outcrops at ataata kuua on the south coast of nuussuaq. the borehole did not reach a change in lithology that could be interpreted as the boundary to the older kingittoq member. locally, the qilakitsoq member is of earliest campanian age (dam et al. 2000; pedersen & nøhr-hansen 2014). the qilakitsoq member has not previously been documented from disko, although a fluvial sandstone of latest cenomanian age overlies the kussinerujuk member at asuk east. it was tentatively correlated to the qilakitsoq member (pedersen & nøhr-hansen 2014), mostly because of its position above the kussinerujuk member. 4.1.1.1 depositional history. the atane formation extends over distances of 150 km n–s and 60 km e–w. it represents deposits from the late albian to the earliest campanian (about 20–25 million years) and ranges from fluvial channel sandstones (skansen member) and complex, estuarine deposits (ravn kløft member) to stacked delta deposits of the kingittoq and qilakitsoq members. in the eastern part of the nuussuaq basin, the non-marine delta facies are dominant, while a stronger marine influence is observed in the western outcrops, especially in the qilakitsoq member. the lithostratigraphic overview of dam et al. (2009, fig. 16) shows that a hiatus between the cretaceous and the palaeogene fig. 5 map showing schematically the known extents of the four members of the atane formation of the nuussuaq group. the lines do not indicate member boundaries but knowledge limitations: members can only be assigned in the areas with numbers 1 to 4. the atane formation is not present in western and northern nuussuaq, except in the area labelled 1 and 3. gv01_02_160_05_lml 70°n 69°30'n 55°w 53°w 55°w 53°w 25 km 70°30'n ice cap west greenland basalt group nuussuaq group precambrian basement 4 qilakitsoq member 3 kingittoq member 2 skansen member 1 ravn kløft member 2 1,3 3 3 4 4 https://doi.org/10.34194/geusb.v57.8361 http://www.geusbulletin.org/ larsen et al. 2024: geus bulletin 57. 8361. https://doi.org/10.34194/geusb.v57.8361 7 of 27 geusbulletin.org sedimentary formations has the longest duration close to the basin boundary fault (eastern disko and eastern nuussuaq), whereas the succession in western nuussuaq is almost continuous. a succession ranging from the lower turonian to the upper coniacian, such as that in cores fp94-4-5 and fp94-4-6, has not previously been documented in the atane formation (e.g. pedersen & nøhr-hansen 2014, fig. 16). 4.1.2 eqalulik formation this formation comprises dark grey to black, marine, synvolcanic mudstones with layers of volcanic ash (dam et al. 2009). it is commonly poorly exposed due to talus deposits from steep cliffs of the overlying hyaloclastite breccias. lateral correlations are rarely possible, except for a distinctive and widespread layer of graphite andesite tuff erupted from the ilugissoq volcano of the asuk member of the vaigat formation in central nuussuaq (fig. 6; a.k. pedersen & larsen 2006; pedersen et al. 2017). dinoflagellate cysts in the eqalulik formation indicate a late danian–early selandian age (nøhr-hansen et al. 2002), which is consistent with radiometric ages of the vaigat formation (storey et al. 1998). the eqalulik formation was deposited during a period of rapid subsidence, at least partly due to loading by the volcanic rocks; water depths of up to c. 600 m are demonstrated by the height of hyaloclastite foresets (pedersen et al. 2017, figs 85, 87). along the north-east coast of disko, the eastern delimitation of the formation is uncertain because it is almost totally unexposed. 4.1.3 atanikerluk formation the synvolcanic, lacustrine sandstones and mudstones in eastern disko and south-eastern nuussuaq constitute the atanikerluk formation (koch 1959; pedersen et al. 1998; dam et al. 2009, figs 127–131). the formation comprises fluvial sandstones (the akunneq member) and lacustrine mudstones grading up into sandstones: the naujât, pingu, umiussat and assoq members. marine dinoflagellate cysts occur locally in the assoq member in southern disko, which suggests intermittent contacts between the lake and marine waters to the south. the atanikerluk formation is bounded to the west by the vaigat formation (ordlingassoq member) and the maligât formation (lower rinks dal member). its extent to the east and south is unknown, and it is possible that the lake was bounded to the east by the basin boundary fault (pedersen et al. 2018, fig. 16). 4.2 igneous rocks 4.2.1 vaigat formation the volcanic rocks of the vaigat formation consist mainly of picrites, i.e. relatively primitive, magnesium-rich rocks that were erupted directly from the melt accumulation areas in the asthenospheric mantle. particularly during the second volcanic episode, some magma batches stalled in the crust and became contaminated with carbon-bearing sediments before eruption. such crustally contaminated rocks range from magnesium-poor picrites through basalts and magnesian basaltic andesites to magnesian andesites and now form five separate, minor members intercalated with the major, uncontaminated members (pedersen et al. 2017). the uncontaminated picrites never reached as far east as the qullissat area; those of the second volcanic episode (naujánguit member) stopped c. 10 km north-west of qullissat, whereas those of the third volcanic episode (ordlingassoq member) stopped just above the site of the town (fig. 6). only the basalt flows of the maligât formation later overran the whole area (fig. 3). 4.2.1.1 asuk member. the asuk member contains the most crustally contaminated rocks in the vaigat formation, with up to 40% sediment component, and is famous for its contents of native iron and graphite (steenstrup 1874, 1875; larsen & pedersen 2009; pedersen et al. 2017). the asuk member was erupted from at least three different centres on nuussuaq and two on disko (fig. 6) close to the end of the second volcanic episode of the vaigat formation. fig. 6 map showing the distribution of the individual units of the asuk member. ce: inferred central eruption area for the asuk member on disko. the large, light grey circular area is the extent of the graphite andesite tuff layer erupted from the ilugissoq volcano. the green areas show exposures of the vaigat formation, and the green lines are the eastern and south-eastern depositional boundaries of the large naujánguit and ordlingassoq members of the vaigat formation. modified from a.k. pedersen et al. (2017, fig 10). ? ? ? ? vaigat st or da l 25 km maligâthareøen 70°30’n 70°n 55°w 53°w 52°w niiortuut unit ilugissoq volcano qullissat tuff range individual units of the asuk member saviaqqat unit asuk mb on disko qullissaaqqat ordlingassoq mb naujánguit mb agatdalen n o ce https://doi.org/10.34194/geusb.v57.8361 http://www.geusbulletin.org/ larsen et al. 2024: geus bulletin 57. 8361. https://doi.org/10.34194/geusb.v57.8361 8 of 27 geusbulletin.org 4.2.1.2 the nunngarut igneous body. the igneous body in the sediments between qullissat and qullissaaqqat is here called the nunngarut igneous body (later: lava flow). it forms an apparently near-horizontal sheet that can be followed along the coast for c. 4.7 km with interruptions due to landslides. where it is in place, its upper boundary is situated just below 300 m altitude (figs 2 and 3), but neither the upper nor the lower boundary are exposed. the body has recently been investigated by magnetic and multispectral methods, by which a number of more or less slipped blocks were identified beside some stable areas (jackisch et al. 2022). the nunngarut igneous body is not present north-west of qullissat and thus occurs about 12 km south-east of the area with lava flows of the asuk member. it even occurs south-east of the mapped south-eastern boundary of the entire vaigat formation (figs 1 and 6). compositionally, the nunngarut igneous body is a magnesian basaltic andesite with 53–56 wt% sio2 and 7.2–8.0 wt% mgo (calculated volatile-free), with highest sio2 in the upper part (olshefsky et al. 1995 and analysis from the exposure at nunngarut in pedersen et al. 2017, table 5). the body is closely similar to the asuk member with respect to mineralogy and composition, and it has tentatively been regarded as a sill belonging to the member. the fp94-4-5 core has for the first time allowed us to study the upper and lower contacts of the body. 5 material and methods data for the fp94-4-5 and fp94-4-6 cores are given in table 1. also included are data for five short sections (m-1 to q-3, figs 2 and 3) measured in the coastal cliffs between qullissaaqqat and the abandoned coal mine and used here for correlation purposes. sections q-1 to q-3 were published by pedersen et al. (2006). when the locations of the drill sites given in olshefsky et al. (1995) were checked in the new high-resolution digital elevation model (styrelsen for dataforsyning og infrastruktur 2023), the given altitudes above sea level of 358 and 86 m for holes fp94-4-5 and fp94-4-6, respectively, did not fit the elevation model and were replaced with our measured altitudes of 315 and 67 m (table 1). the universal transverse mercator (utm) coordinates appeared to be correct. core fp94-4-5 starts at 30.5 m downhole depth and core fp94-4-6 starts at 28.55 m downhole depth because the overlying parts consisted of loose overburden of which no cores are preserved. the cores have a diameter of c. 3.6 cm, and they are not slabbed. the preservation state of the rocks is excellent, and the stratification in the laminated sediments appears to be horizontal. the sedimentological logs show the estimated grain sizes and the sedimentary structures, which form the basis for the identification of sedimentary facies and the interpretation of the depositional environment when combined with the palynological data. the samples for palynological examination were treated in the lab as described in g.k. pedersen et al. (2023). all palynological slides and (if available) organic residues are stored at geus. total carbon (tc), total organic carbon (toc) and total sulphur (ts) analyses, rock-eval type pyrolysis and vitrinite reflectance measurements were carried out following procedures detailed by andrews et al. (2022). when intervals and metres are mentioned in the core descriptions and discussions, they are downhole depths unless altitudes above sea level (m a.s.l.) are specified. 6 results figure 7 shows overview logs of the two cores in their actual position relative to each other, i.e. relative to the sea level. the nunngarut igneous body (the drilling target) with 138 m thickness takes up most of the length of core fp94-4-5. the entire sedimentary succession in core fp94-4-6 and the lower part (270–205 m downhole depth) of the sedimentary succession in core fp94-4-5 are interpreted as deltaic deposits of the atane formation. based on the palynological results and the sedimentary facies, the overlying succession in core fp94-4-5 (205–190 m downhole depth) table 1 data for investigated drill cores and surface sections in the qullissat area. core or section coordinates utm zone 22 length altitudes, m a.s.l. deg. w deg. n easting northing m top bottom fp94-4-5 53° 00.04’ 70° 02.854’ 423818 7772428 270.5 315 44.5 fp94-4-6 52° 58.38’ 70° 03.147’ 424891 7772939 143.26 67 –76.26 q-1 52° 58.02’ 70° 03.236’ 425121 7773097 29 29 0 q-2 52° 58.07’ 70° 03.266’ 425094 7773152 26 27 1 q-3 52° 58.32’ 70° 03.449’ 424931 7773499 53 53 0 m-1 52° 58.95’ 70° 03.856’ 424568 7774268 27 42 15 m-2 52° 59.12’ 70° 03.970’ 424469 7774483 22 22 0 cores: utm coordinates are as reported in olshefsky et al. (1995); geographical coordinates and altitudes were calculated/measured at geus. the cores are stored in geus’s core repository, copenhagen. surface sections q1 to m2: all coordinates were measured/calculated at geus. m a.s.l. = metres above sea level. https://doi.org/10.34194/geusb.v57.8361 http://www.geusbulletin.org/ larsen et al. 2024: geus bulletin 57. 8361. https://doi.org/10.34194/geusb.v57.8361 9 of 27 geusbulletin.org is referred to the paleocene eqalulik formation. the sedimentary succession above the nunngarut igneous body is referred to the paleocene atanikerluk formation. the horizontal distance between boreholes fp94-4-5 and fp94-4-6 is c. 1190 m (table 1). vertically, the cores appear to be almost in stratigraphic continuity, as core 5 ends at 45.5 m a.s.l. and core 6 begins at 38.5 m a.s.l. (fig. 7 and table 1), and the sediments are near-horizontal. nonetheless, the sedimentary successions and the biostratigraphy in the two cores may be correlated as shown below, suggesting the presence of a fault between the two core sites (fig. 7). 6.1 palynology and ages a total of 24 samples from fp94-4-5 (ggu no. 393016) and 20 samples from fp94-4-6 (ggu no. 393017) have been processed, and the palynomorphs have been studied. the results are presented in range charts and photographs of stratigraphically important species in figs 8–11. the samples are located in the detailed core logs (fig. 12). marine dinoflagellate cyst species are common, although their abundances in many samples are low. several markers (fo: first occurrence and lo: last occurrence of species) also occur in both cores as indicated in the range charts. some samples containing only spores and pollen probably represent non-marine environments. 6.1.1 core fp94-4-5 samples 101 and 108 are barren of dinoflagellate cyst and miospores. samples 104, 105 and 106 contain a few miospores. samples 102 (253 m) and 103 (251 m) contain a few well-preserved specimens of heterosphaeridium difficile and few, not very age-diagnostic dinoflagellate cyst species and the pollen rugubivesiculites rugosus, together indicating an age not older than turonian. fig. 7 simplified overview logs of the investigated cores fp94-4-5 and fp94-4-6 showing their position relative to sea level (scale at left). downhole depths are shown along each core. correlation of the cretaceous sediments in the two cores is based on their sedimentological logs (see below). a characteristic correlatable drowning surface is shown in blue. the surface is situated c. 60 m above sea level in core fp94-4-5 and c. 30 m below sea level in core fp94-4-6, indicating a vertical displacement between the drill sites of c. 90 m along a fault. the fault is schematically indicated on the drawing; its precise position and orientation are uncertain. 150 100 95 sea level 0 50 fp94-4-6 fp94-4-5 overburden (no core) overburden (no core) paleocene sediments paleocene sediments cretaceous sediments cretaceous sediments top breccia nunngarut igneous body (lava flow) 200 190 270 250 255 9 0 m 150 100 50 0 depth in core (m) distance c. 1 km (fig. 1b) drowning surface fig. 17 depth in core (m)m et re s ab ov e se a le ve l 100 50 0 –50 –100 150 200 250 300 https://doi.org/10.34194/geusb.v57.8361 http://www.geusbulletin.org/ larsen et al. 2024: geus bulletin 57. 8361. https://doi.org/10.34194/geusb.v57.8361 10 of 27 geusbulletin.org samples 107 (216 m) and 109 (206 m) also contain heterosphaeridium difficile and rugubivesiculites rugosus besides isabelidinium and chatangiella specimens. sample 109 also contains a poorly preserved raphidodinium fucatum specimen indicating a post middle turonian age, whereas both samples contain chatangiella mcintyrei which may indicate an early coniacian age according to nøhr-hansen (1996) and pedersen & nøhr-hansen (2014). samples 110 (201 m) to 118 (190 m) are of early paleocene age, possibly danian. sample 111 (200 m) contains only one in situ dinoflagellate cyst species but a fair number of specimens (18); the species is similar to dinocyst sp. 3, which was recorded from the eqalulik formation on the north coast of nuussuaq by nøhr-hansen et al. (2002). samples 110 (201 m) to 117 (191 m) contain a few reworked upper cretaceous dinoflagellate cyst specimens (isabelidinium/chatangiella), the pollen rugubivesiculites rugosus and, in sample 111 (200 m), the maastrichtian marker pollen species wodehouseia spinata. sample 118 (190 m, at the contact to the nunngarut igneous body) only contains coal fragments. sample 119 (51 m) contains some small round thickwalled brown bodies. samples 120 (47 m) and 121 (46 m) contain few reworked upper cretaceous dinoflagellate cysts and the pollen rugubivesiculites rugosus. sample 122 (34.2 m) contains three specimens of the dinoflagellate cyst taurodinium granulatum (formerly classified as an acritarch). sample 100 (32 m; taken before the systematic sampling) contains some taurodinium granulatum and some specimens of the freshwater alga pediastrum. sample 123 (31 m) contains no dinoflagellate cysts but is dominated by bisaccate pollen. piasecki et al. (1992) first reported taurodinium granulatum as gen et sp. indet. from tuapaat in south-eastern disko in a sand and mudstone succession in the rinks dal member. storey et al. (1998) dated the rinks dal member to c. 61 ma (selandian). it is therefore assumed that the presence of the species in samples 122 and 100 represents a selandian age. for further record of taurodinium granulatum, see fensome et al. (2016, p 72). fig. 8 range chart of dinoflagellate cysts, spores and pollen (sp), acritarchs, fungi (fu) and freshwater algae (al) in core fp94-4-5. biostratigraphically important palyno-events are shown to the right. fo: first occurrence. lo: last occurrence. d ep th 50m 75m 100m 125m 150m 175m 200m 225m fp94-4-5 393016 c hr on ost ra ti gr ap hy pa le oc en e la te c re ta ce ou s pe ri od /e po ch ?s el an di an d an ia n coniacian ?l at e tu ro ni an ?c on ia ci an a ge li th ost ra ti gr ap hy a ta ni ke rl uk v ai ga t eq al ul ik a ta ne fo rm at io n sa m pl es 31.00 123 32.00 100 32.50 122 46.00 121 47.00 120 51.00 119 190.00 118 191.00 117 192.00 116 194.00 115 196.00 114 197.50 113 199.00 112 200.00 111 201.00 110 206.00 109 207.00 108 216.00 107 221.00 106 232.50 105 240.00 104 251.00 103 253.00 102 256.00 101 c ha ta ng ie lla s pp . c irc ul od in iu m s pp . c le is to sp ha er id iu m a ci cu la re h et er os ph ae rid iu m d i� ci le o lig os ph ae rid iu m c om pl ex o lig os ph ae rid iu m p ul ch er rim um pa la eo pe rid in iu m c re ta ce um c ha ta ng ie lla m ci nt yr ei o do nt oc hi tin a op er cu la ta c le is to sp ha er id iu m sp p. is ab el id in iu m c oo ks on ia e is ab el id in iu m s pp . pa la eo hy st ric ho ph or a ch ei t ra ph id od in iu m fu ca tu m sp in ife rit es sp p. d in oc ys t s p 3 h n h 2 00 2 18 1 2 3 3 2 1? 1 2? 6 1? 1 1 3 3 1 1 1 1? 4 1? 3 1 1? 1? 1 1 2? 4 4 dino�agellate cysts reworked c ha ta ng ie lla s pp . c le is to sp ha er id iu m a ci cu la re h et er os ph ae rid iu m d i� ci le c le is to sp ha er id iu m s pp . is ab el id in iu m s pp . sp in ife rit es s pp . is ab el id in iu m c f. cr et ac eu m o lig os ph ae rid iu m s pp . 1 1? 1 1 1 1 1 1 1 1 1 3 1 1 1? dino�agellate cysts ru gu bi ve sic ul ite s r ug os us a pp en di ci sp or ite s i ns ig ni s in ap er tu ro po lle ni te s h ia tu s 2? 6 1 7 3 9 1 3 7 2 sp reworked ru gu bi ve sic ul ite s r ug os us a pp en di ci sp or ite s i ns ig ni s a qu ila po lle ni te s s pp . w od eh ou se ia sp in at a ba lm ei sp or ite s g le ne lg en sis 3 1 1 1 2 3 1 3 1 1 1 1 1 spores & pollen a cr ita rc h sp p. pa ra le ca ni el la in de nt at a ta ur od in iu m g ra nu la tu m 7 1 1 1 acritarchs fu ng al s pp . 1 1 1 1 1 fu pe di as tr um s pp . 4 1 al events 32.50 taurodinium granulatum 200.00 dinocyst sp 3 hnh 2002 206.00 raphidodinium fucatum 216.00 chatangiella mcintyrei 253.00 heterosphaeridium di�cile, rugubivesiculites rugosus fo https://doi.org/10.34194/geusb.v57.8361 http://www.geusbulletin.org/ larsen et al. 2024: geus bulletin 57. 8361. https://doi.org/10.34194/geusb.v57.8361 11 of 27 geusbulletin.org fig. 9 palynomorphs from core fp94-4-5. same magnification for all figures; scale bar in aa = 20 μm. each figure represents species name, sample no., slide no., england finder coordinate and photo no. a: taurodinium granulatum, sample 393016-100, slide 2, w52-3, photo 2709. b: taurodinium granulatum, sample 393016-100, slide 3, p28-4, photo 2711. c: taurodinium granulatum, sample 393016-100, slide 2, w54-1, photo 2713. d: pediastrum sp., sample 393016-100, slide 2, f41-1, photo 2710. e: taurodinium granulatum, sample 393016-122, slide 2, h57-1, photo 2708. f: dinocyst sp. 3 nøhr-hansen et al. (2002), sample 393016-111, slide 6, l19, photo 2582. g: dinocyst sp. 3 nøhr-hansen et al. (2002), sample 393016-111, slide 6, l30-4, photo 2586. h: dinocyst sp. 3 nøhr-hansen et al. (2002), sample 393016-111, slide 6, f42-1, photo 2595. i: dinocyst sp. 3 nøhr-hansen et al. (2002), sample 393016-111, slide 6, z20-2, photo 2583. j: dinocyst sp. 3 nøhr-hansen et al. (2002), sample 393016-111, slide 6, w26-2, photo 2585. k: dinocyst sp. 3 nøhr-hansen et al. (2002), sample 393016-111, slide 6, x38-1, photo 2591. l: dinocyst sp. 3 nøhr-hansen et al. (2002), sample 393016-111, slide 6, f421, photo 2592. m: dinocyst sp. 3 nøhr-hansen et al. (2002), sample 393016-111, slide 6, x32, photo 2588. n: dinocyst sp. 3 nøhr-hansen et al. (2002), sample 393016-111, slide 6, p42-2, photo 2699. o: dinocyst sp. 3 nøhr-hansen et al. (2002), sample 393016-111, slide 6, o42-4, photo 2581. p: dinocyst sp. 3 nøhr-hansen et al. (2002), sample 393016-111, slide 6, g23-3, photo 2584. q: aquilapollenites sp. reworked, sample 393016-111, slide 4, u31-1, photo 2704. r: wodehousia sp. reworked, sample 393016-111, slide 6, h20-4, photo 2701. s: isabelidinium cf. cretaceum reworked, sample 393016-112, slide 6, x37-4, photo 2706. t: isabelidinium cf. cretaceum reworked, sample 393016-111, slide 6, s25-4, photo 2702. u: chatangiella mcintyrei, reworked, sample 393016-111, slide 6, q29-1, photo 2700. v: chatangiella mcintyrei, sample 393016-107, slide 6, q21-1, photo 2691. w: heterosphaeridium difficile, sample 393016-107, slide 6, l49-1, photo 2693. x: rugubivesiculites rugosus, sample 393016-106, slide 5, l46-2, photo 2698. y: heterosphaeridium difficile, sample 393016-102, slide 5, g45-4, photo 2688. z: raphidodinium fucatum, sample 393016-109, slide 6, q27-4, photo 2697. aa: raphidodinium fucatum, sample 393016-109, slide 4, g28-1, photo 2698. ab: heterosphaeridium difficile, sample 393016-102, slide 5, x39, photo 2687. https://doi.org/10.34194/geusb.v57.8361 http://www.geusbulletin.org/ larsen et al. 2024: geus bulletin 57. 8361. https://doi.org/10.34194/geusb.v57.8361 12 of 27 geusbulletin.org in conclusion, the atane formation in fp94-4-5 is interpreted to be of late turonian to coniacian age. the sediments at 205–201 m are assumed to be danian, and those at 201–190 m are of danian age and are referred to the eqalulik formation. the sediments above the nunngarut igneous body at 51–31 m are of selandian age and are referred to the atanikerluk formation. 6.1.2 core fp94-4-6 samples 1, 8, 10, 11 and 16 are barren of dinoflagellate cysts, and samples 3, 4, 5, 17 and 18 are nearly so. sample 2 (134 m) contains trithyrodinium suspectum (indicating a late cenomanian or younger age), questionable heterosphaeridium difficile and chatangiella granulifera specimens (indicating a turonian or younger age). the dinoflagellate cyst assemblage is dominated by a species tentatively described as chatangiella aff. multispinosa; the species was originally described from the albian to cenomanian of western australia by cookson & eisenack (1970) and later reported from the santonian on antartica as chatangiella cf. multispinosa by keating (1992). the species heterosphaeridium difficile is present in five samples from sample 2 (134 m) to sample 20 (29 m), with an acme in sample 12 (70.5 m). the acme of heterosphaeridium difficile together with ?raphidodinium fucatum and ?chatangiella mcintyrei indicates a coniacian age according to nøhr-hansen (1996), pedersen & nøhr-hansen (2014) and bailey & biostrat (2021). the possible last occurrence (lo) of chatangiella mcintyrei at sample 15 (56.3 m) and the presence of heterosphaeridium difficile in the uppermost sample 20 (29 m) indicate a minimum late coniacian or early santonian age according to bailey & biostrat (2021). the pollen rugubivesiculites rugosus is present in all the 20 studied samples. in conclusion, the atane formation in fp94-4-6 is of turonian (143–70.5 m) and coniacian (70.5–29 m) age. 6.1.3 palynology of coal seams b–c in the coastal cliff at qullissaaqqat palynological samples from coal seams b and c from coastal section q-3 (see fig. 17) are all dominated by spores, primarily from ferns but also from bryophytes, while pollen from gymnosperms and angiosperms are rare, the latter two groups making up 7–26% of the spore-pollen content in coal seam b and 9% in coal seam c. this suggests that both coal seams represent relatively open mires that were not dominated by trees or bushes. this differs somewhat from the interpretation made by pedersen et al. (2006), who suggested that trees and large bushed were present in the mire flora, based on the coal petrographic dominance of huminite and occasionally inertinite, but with a slightly higher proportion of herbaceous plants in coal seam c. fig. 10 range chart of dinoflagellate cysts, spores and pollen, and acritarchs (ac) in core fp94-4-6. biostratigraphically important palyno-events are shown to the right. fo: first occurrence. lo: last occurrence d ep th 40m 60m 80m 100m 120m 140m c hr on ost ra ti gr ap hy la te c re ta ce ou s pe ri od /e po ch ?c on ia ci an tu ro ni a n a ge sa m pl es 29.00 20 32.30 19 39.20 18 40.20 17 49.80 16 55.00 14 56.30 15 64.80 13 70.50 12 78.40 11 90.00 10 92.30 9 95.00 8 109.30 7 111.90 6 120.80 5 124.80 4 127.30 3 134.00 2 143.00 1 fp 94-4-6 393017 c ha ta ng ie lla g ra nu lif er a c ha ta ng ie lla a �. m ul tis pi no sa c irc ul od in iu m d is tin ct um c le is to sp ha er id iu m a ci cu la re d es m oc ys ta p le kt a fl or en tin ia s pp . h et er os ph ae rid iu m d i� ci le o do nt oc hi tin a an ca la o do nt oc hi tin a op er cu la ta o lig os ph ae rid iu m c om pl ex pa la eo pe rid in iu m c re ta ce um sp in ife rit es s pp . tr ith yr od in iu m su sp ec tu m is ab el id in iu m s pp . c ha ta ng ie lla s pp . c irc ul od in iu m s pp . c or on ife ra st rio la ta h ys tr ic ho di ni um p ul ch ru m o lig os ph ae rid iu m p ul ch er rim um su rc ul os ph ae rid iu m lo ng ifu rc at um c ha ta ng ie lla m ci nt yr ei is ab el id in iu m c oo ks on ia e is ab el id in iu m m ag nu m ra ph id od in iu m fu ca tu m ex oc ho sp ha er id iu m s pp . c le is to sp ha er id iu m s pp . 1 1? 1 1? 1? 1 1? 4 1 1 1? 2 1? 2 1 1 1 1 1 1 1? 1 23 1 1? 1? 1? 1? 5 1 1 1 3 1 1 1 1 2 1 4 5 1 1 1 7 16 1 1 10 1 1 3 1 1 1? 69 12 6 1 1 1? 1? 1 3 1 3 1 dino�agellate cysts a cr ita rc h sp p. 3 ac a pp en di ci sp or ite s i ns ig ni s ru gu bi ve sic ul ite s r ug os us ru gu bi ve sic ul at es re du ct us st ov er isp or ite s l un ar is h oe gi sp or is sp p. d ic ha st op ol le ni te s d un ve ga ne ns is 1 13 1 7 3 5 1 8 1 8 1 1 15 1? 26 4 1 12 1 5? 15 1? 1 1 4 5 1 1 1 17 1? events 56.30 70.5070.50 chatangiella mcintyrei raphidodinium fucatum, ?chatangiella mcintyrei heterosphaeridium di�cile acme 134.00134.00 heterosphaeridium di�cile, trithyrodinium suspectum, rugubivesiculites rugosus chatangiella a�. multispinosa acme lo fo spores & pollen https://doi.org/10.34194/geusb.v57.8361 http://www.geusbulletin.org/ larsen et al. 2024: geus bulletin 57. 8361. https://doi.org/10.34194/geusb.v57.8361 13 of 27 geusbulletin.org fig. 11 palynomorphs from core fp94-4-6. same magnification for all figures; scale bar in t = 20 μm. each figure represents species name, sample no., slide no., england finder coordinate and photo no. a: heterosphaeridium difficile, sample 393017-15, slide 5, e50-3, multifocus. b: chatangiella mcintyrei, sample 393017-15, slide 4, g48-4, photo 2741. c: raphidodinium fucatum, sample 393017-12, slide 4, w34-4, photo 2740. d: chatangiella mcintyrei, sample 393017-12, slide 5, s19-1, photo 2739. e: chatangiella mcintyrei, sample 393017-12, slide 4, t50-4-1, photo 2738. f: heterosphaeridium difficile, sample 393017-12, slide 2, c31-3, photo 2737. g: heterosphaeridium difficile, sample 393017-9, slide 5, j22-4, photo 2735. h: rugubivesiculites rugosus, sample 393017-9, slide 4, h29-4, photo 2736. i: heterosphaeridium difficile, sample 393017-9, slide 5, e53-2, photo 2733. j: heterosphaeridium difficile, sample 393017-6, slide 4, a32-1, photo 2732. k: heterosphaeridium difficile, sample 393017-6, slide 4, l55-4, photo 2731. l: spiniferites sp., sample 393017-6, slide 4, b52-3, photo 2729. m: appendicisporites insignis, sample 393017-3, slide 4, w34-1, photo 2728. n: ? heterosphaeridium difficile, sample 393017-2, slide 4, s34-2, photo 2735. o: trithyrodinium suspectum, sample 393017-2, slide 5, c28-3, photo 2724. p: chatangiella cf. granulifera, sample 3930172, slide 4, x26-1, photo 2720. q: chatangiella aff. multispinosa, sample 393017-2, slide 4, h32.4, photo 2719. r: chatangiella aff. multispinosa, sample 393017-2, slide 4, q47-3, photo 2718. s: chatangiella aff. multispinosa, sample 393017-2, slide 4, p19-2, photo 2716. t: chatangiella aff. multispinosa, sample 393017-2, slide 4, p19-2, photo 2715. https://doi.org/10.34194/geusb.v57.8361 http://www.geusbulletin.org/ larsen et al. 2024: geus bulletin 57. 8361. https://doi.org/10.34194/geusb.v57.8361 14 of 27 geusbulletin.org 6.2 vitrinite reflectance and organic geochemistry three core samples were analysed for vitrinite reflectance and organic geochemistry: a carbonaceous mudstone 10 cm below the base of the nunngarut igneous body, a coal seam 2.3 m above the body and a carbonaceous mudstone c. 1 m above the body. the last sample is from a short and very broken core from hole 94-4-4, situated c. 2 km north of hole 94-4-5 (olshefsky et al. 1995). the results are presented in table 2. the organic matter in the sample from below the base of the igneous body has a very high vitrinite reflectance (ro) of 1.67% ro and a similarly high tmax of 600°c. except for the deep wells gro#3 and umiivik-1, these values are far higher than for other investigated samples in the nuussuaq basin which have ro ≤ 0.57% and tmax ≤ 450°c (shekhar et al. 1982; bojesen-koefoed et al. 2001; pedersen et al. 2006). even in the gro#3 well, values of 1.67% ro are only attained at depths of 2000 m, and tmax does not exceed 575°c (bojesen-koefoed et al. 1997). comparable values are found only in the lower part of the umiivik-1 core (2–4% ro and 470–580°c tmax), where they are attributed to the heating effect of several thick sill intrusions at this level (dam et al. 1998a). the organic matter in the two samples above the igneous body has vitrinite reflectances of 0.35% ro and 0.50% ro and tmax of 414°c and 428°c (table 2); these values are within the known and normal ranges for fig. 12 sedimentological logs of cores fp94-4-5 and fp94-4-6. m 60 70 80 90 30 40 50 100 110 130 140 143.26 120 vf silt sand grclay f m c vc ja 20 19 18 17 16 14 15 13 12 11 10 9 8 7 6 5 4 3 2 1 a ta ne f m , q ila ki ts oq m em be r tu ro ni an c on ia ci an ? ? ? ? ? u pp er c re ta ce ou s a ta ni ke rl uk f or m at io n eq al ul ik f m a ta ne f m , q ila ki ts oq m em be r pa le oc en e u pp er c re ta ce ou s tu ro ni an va ig at f m ja,gy ? ? vf silt sand grclay f m c vc lithology sand with coal debris coal mudstone silt-streaks sand-streaks sandstone cementation jarosite, gypsum pyrite ja,gy pebbly sandstone lava sedimentary structures structureless cross-lamination wave-ripple cross-lamination hummocky cross-stratification soft sediment deformation cross-bedding trough cross-bedding trace fossils depositional environments chondrites bioturbation planolites isp. ophiomorpha isp. escape trace teichichnus isp. diplocraterion parallelum parallel lamination fossils plant remains rootlets coalified wood marine dinoflagellate cysts fluvial channel delta plain, floodplain delta front and interdistributary bay shoreface deep water marine lacustrine c on ia ci an fp 94-4-5 123 122 121 120 119 118 116 115 114 113 112 111 110 109 108 107 106 105 104 102 103 101 fp 94-4-6 270 260 240 220 210 190 50 40 30 m 230 https://doi.org/10.34194/geusb.v57.8361 http://www.geusbulletin.org/ larsen et al. 2024: geus bulletin 57. 8361. https://doi.org/10.34194/geusb.v57.8361 15 of 27 geusbulletin.org samples from the nuussuaq basin stated earlier. these values translate into a maximum burial temperature of 50–80°c (barker & pawlewicz 1994), which, allowing for a higher-than-normal geothermal gradient related to volcanic activity, suggests a depth of burial of less than about 2 km. this is, in turn, in keeping with data from the gro#3 well (bojesen-koefoed et al. 1997). the sediment below the igneous body has obviously been heated by the emplacement of the body. the heating effect reaches 4 m into the sediment, as judged from the slightly increased magnetic susceptibility of the sediment in this interval (olshefsky et al. 1995, their appendix ii), probably caused by re-magnetisation of magnetite grains during the strong heating. the sediments above the body do not show any trace of being heated, and although the sample distances to the body are greater than for the sample below the base, signs of heating of the sediment on the top should be detectable if the body is a sill, as in the umiivik-1 well. the absence of heating is simply explicable if the body is a lava flow. 6.3 atane formation the sedimentological logs of the two cores (fig. 12) show a range of sedimentary facies from cross-bedded sandstone, through heterolithic sandstone and sand-streaked mudstone, to mudstone and coal beds, illustrated by core photographs in fig. 13. they are interpreted as representing four major depositional environments indicated with colour codes in fig. 12. 6.3.1 fluvial and distributary channel sandstones cross-bedded, mediumto coarse-grained sandstones, locally with comminuted plant debris, and without marine trace fossils, are typically 2–10 m thick (average c. 4 m) and weakly upwards fining (fig. 12). examples of these sandstones occur at 141–135 m and 109–98 m in fp944-6 and at 256–270 m in fp94-4-5. they are interpreted as deposited in fluvial and distributary channels. the channel sandstones constitute close to one-third of the atane formation in the two cores. thick sandstone units may represent two phases of fluvial deposition (fig. 12; section q-3 at 34–25 m, see fig. 17). upwards fining sandstones are interpreted as deposited in minor channels. in the slim cores, intervals of weakly cemented sandstone often have low recovery and occur as loose sand in the core boxes. the base of the fluvial sandstones may truncate the top of delta front successions. the fluvial sandstones are typically overlain by delta-plain deposits. 6.3.2 delta plain deposits mudstones with plant debris interbedded with black coal-bearing mudstones of varying thicknesses and thin sandstone beds characterise the delta plain facies association (fig. 13b, c). the delta plain units are 1–9 m thick (average c. 4 m) and constitute about one-third of the atane formation in the two cores. thin, upwards coarsening units, associated with coal-rich beds, may have formed during mouth-bar progradation in interdistributary bays. examples of delta plain successions occur at 129–124 m and 96.7–94.6 in fp94-4-6 and 209.8–206.5 m in fp94-4-5. chemical analyses show that the black mudstones typically contain less than 50% of organic carbon and thus, strictly speaking, are not coal beds (table 2). modern analogues for ancient coals are peat-accumulating mires, which have low contents of clastic material and sulphur, and where peat accumulates in beds that are at least 10 m thick. such conditions occur in floating mires, low-lying mires and raised mires (haszeldine 1989). layers of peat are often reduced to 10% of their original thickness during compaction to coal (ryer & langer 1980; ryer 1981). thick coal beds therefore represent fairly long periods of time and commonly appear to be laterally continuous. 6.3.3 transgressive shoreface sandstones thin beds 0.1–0.5 m thick, of fineto medium-grained, wellsorted sandstone with an erosive basis and sometimes with marine trace fossils, separate delta plain deposits (below) from delta front deposits (above). the transgressive shoreface sandstones are rare and constitute c. 1% table 2 vitrinite reflectance (ro) and screening data for organic-rich samples below and above the nunngarut lava flow. core fp94-4-5 fp94-4-4 fp94-4-5 ggu-no of core 393016 393015 393016 depth in hole (m) 190.4 48.3 50.5 distance to lava (m) 0.1 (to base) c. 1 (to top) 2.3 (to top) material carbonacous mudstone carbonaeous mudstone coal ro (%) 1.67 0.50 0.36 sd 0.14 0.04 0.02 n (no of points) 128 100 100 toc (wt%) 1.61 6.77 43.54 tc (wt%) 1.84 6.77 45.92 ts (wt%) 0.70 0.13 2.07 tmax (°c) 600 428 414 s1 (mg hc/g rock) 0.17 0.12 0.65 s2 (mg hc/g rock) 0.90 8.25 51.97 s3 (mg co2/g rock) 2.47 20.91 hi 56 122 119 oi 153 48 pi 0.16 0.01 0.01 pc 0.09 0.69 4.37 sd: 1 s standard deviation. s1: free hydrocarbons (rock-eval type pyrolysis). s2: pyrolytic hydrocarbons (rock-eval type pyrolysis). s3: co2 from rock-eval type pyrolysis. hi: hydrogen index ((100*s2)/toc). oi: oxygen index ((100*s3)/toc). pi: production index (s1/(s1+s2)). pc: pyrolytic carbon (0.083*(s1+s2)). toc: total organic carbon. tc: total carbon. ts: total sulphur. https://doi.org/10.34194/geusb.v57.8361 http://www.geusbulletin.org/ larsen et al. 2024: geus bulletin 57. 8361. https://doi.org/10.34194/geusb.v57.8361 16 of 27 geusbulletin.org of the atane formation in the cores. examples occur at 48 m and 119 m in fp94-4-6 and at 254 m in fp94-4-5 (fig. 12). transgressive sandstone beds mark longer breaks in sediment accumulation: at qullissaaqqat, such beds are few and thin, which suggests a palaeogeographic position protected from frequent marine reworking. 6.3.4 delta front deposits figure 12 shows that coarsening upwards (cu) successions, up to 11 m thick, occur through the cored interval of the atane formation. most delta front units are thin, on average c. 4 m thick, and constitute close to one-third of the atane formation in the cores. this facies association ranges from dark grey mudstone and silt-streaked mudstone through interbedded sandstone and mudstone (fig. 13f) to sandstone with hummocky cross-stratification (fig. 13 g–i). comminuted plant debris occurs in the sandstones. marine trace fossils are present, but the degree of bioturbation is generally low. marine dinoflagellate cysts are present in many samples from this facies association, but commonly in low numbers. the cu units are interpreted as deposited during delta front progradation. thick cu units, such as 94–84 m and 44–39 m in core fp94-4-6, may be correlated laterally (see fig. 17). in the paatuut area (south coast of nuussuaq), the delta front deposits constitute 35–40% of the qilakitsoq member, and the cu units range in thickness from few metres to c. 20 m (dam et al. 2009, fig. 46). the thinner cu units in the qullissaaqqat area suggest that the entire succession here may represent more proximal parts of the delta. 6.4 eqalulik formation the eqalulik formation was cored in fp94-4-5 at 205– 190 m (fig. 12). the basal 4 m consists of structureless fig. 13 photographs of the atane formation sediments in the cores. core diameter is 36 mm. a: fineto medium-grained, cross-bedded sandstone, locally with coalified plant debris, interpreted as fluvial or distributary channel facies. core fp94-4-5, depth c. 237.70–237.50 m. b: delta plain facies, dark grey mudstone with lenses or thin layers of coal. core fp94-4-5, depth c. 238.70–238.55 m. c: delta plain mudstone erosively overlain by bioturbated muddy sandstone, interpreted as an interdistributary bay deposit. the boundary represents an initial transgression. core fp94-4-6, depth c. 94.64–94.43 m. the bioturbated facies is overlain by a thick deltafront succession at c. 94.05–81.70 m depth, see fig. 12. d: bioturbated muddy sandstone, close-up of the trace fossil taenidium isp.; the thorough mottling obscures many individual burrows. core fp94-4-6, depth c. 94.5 m (see c). e: dark grey mudstone, interpreted as delta front facies. core fp94-4-6, depth c. 49.75–49.60 m. f: sand-streaked mudstone, alternation between thin layers of very fine-grained sand and layers or laminae of silty mudstone, interpreted as delta front facies. note the scarcity of burrows. core fp94-4-5, depth c. 215.10–214.95 m. g–h–i: delta front facies. well-sorted, fine-grained sandstone with thin drapes of silt and coalified plant debris. local precipitation of pyrite. the sandstone has hummocky cross-stratification (g–i) and wave-ripple cross-stratification (g–h). trace fossil are scarce, suggesting rapid deposition and possibly marine reworking. core fp94-4-6, depth c. 90.10–86.35 m. lengths of core pieces are 16–18 cm. ge f iha b c d https://doi.org/10.34194/geusb.v57.8361 http://www.geusbulletin.org/ larsen et al. 2024: geus bulletin 57. 8361. https://doi.org/10.34194/geusb.v57.8361 17 of 27 geusbulletin.org sandstone (fig. 14a), which probably was deposited during a marine transgression. the homogeneous sandstone facies is overlain by silt-streaked black mudstones (fig. 14b). the mudstone locally contains thin streaks of sand-sized light grains in the mudstone matrix (fig. 14c). this texture is tentatively interpreted as deposition from cohesive mudflows in which the matrix strength was sufficient to prevent the coarse sand from settling (talling et al. 2012). the laminated, silt-streaked mudstone (fig. 14d) is interpreted to have been deposited at some distance from the coastline and at depths well below wave-base, as no indications of wave-reworking have been recognised. distinct tuff beds have not been recognised in the formation in core fp94-4-5 but are known from other sections (dam et al. 2009). the uppermost mudstone has a baked zone of at least 10 cm thickness at the contact to the overlying nunngarut igneous body. 6.5 atanikerluk formation the sediment overlying the nunngarut igneous body at 52.0–50.9 m consists of light grey, irregular, altered volcanic clasts 1–5 mm in size in a weakly bedded, dark violet-grey matrix that contains quartz grains (figs 15, 16a). there are scattered, small (1–2 mm) black clasts and red-oxidised patches in the matrix. the clast size decreases upwards. at 50.9–50.3 m, the rock is a siltstone heavily crushed by the drilling but including 15 cm massive coal in the middle. this is overlain by cross-bedded sandstone (50.3–35.4 m), 0.3 m drilling-disturbed volcaniclastic sediment with quartz grains, a small intrusion? (33.8–32.6 m) and 2 m of mudstone (32.6–30.5 m). recovery was moderate above 45 m, and no core was recovered above 30.5 m. the lower part of the sandstone interval comprises medium-grained, cross-bedded sandstone (50.3–46.2 m, fig. 14e–f) with some coal debris. it may represent fluvial fig. 14 photographs of the eqalulik and atanikerluk formation sediments in the cores. a–d: eqalulik formation; e–g: atanikerluk formation; all from core fp94-4-5, diameters 36 mm. a: structureless sandstone with very small mudstone clasts in a mudstone matrix. lower part of the eqalulik formation, depth c. 201.60–201.45 m. b: dark grey mudstone with sharp-based streaks and thin layers of pale, very fine-grained sand and coarse silt. the thickest layers (3–7 mm) are faintly graded. red-brown colour is due to cementation. depth c. 200.55–200.40 m. c: mudstone facies characterised by thin layers of angular, light grey grains in a mudstone matrix, interpreted as deposited from thin mudflows with sufficient matrix strength to transport the sandto granule-sized grains. depth c. 198.95–198.80 m. note the sharp boundary to homogeneous mudstone at c. 198.85 m. d: grey, silt-streaked mudstone, marine deep-water deposit, upper part of the eqalulik formation, depth c. 191.95–191.80 m. e: cross-bedded, grey sandstone with tiny mudstone clasts. atanikerluk formation, depth c. 48.08–47.95 m. f: cross-bedded, dark grey sandstone with numerous tiny mudstone clasts, depth c. 46.45–46.30 m. g: angular and rounded, moderately sorted clasts in a matrix od sand and mud, interpreted as colluvial deposits. depth c. 36.90–36.75 m. ca b ed gf https://doi.org/10.34194/geusb.v57.8361 http://www.geusbulletin.org/ larsen et al. 2024: geus bulletin 57. 8361. https://doi.org/10.34194/geusb.v57.8361 18 of 27 geusbulletin.org deposits. the overlying fine-grained sandstone (46.2–45.3 m) appears to be hummocky cross-stratified suggesting wave-reworking. two metres of intraformational conglomerate with sand and mudstone clasts in a clayey matrix (37.3–35.4 m, fig. 14g) may represent mudflow deposits (colluvial deposits). the small intrusion? of fine-grained, aphyric basalt is emplaced between the sandstones and overlying mudstones. it is at least 1.4 m thick, but no contacts are preserved. the mudstones range from dark brown, structureless clay with thin beds of yellow sand to greyish silty mudstone with thin streaks of silt and very fine-grained sand. the occurrence of the dinoflagellate cyst taurodinium granulatum (see section on palynology) suggests near-contemporaneity with the lower part (rinks dal member) of the maligât formation (fig. 4). 6.6 the nunngarut igneous body the complete body was penetrated in hole fp94-4-5 from 52 to 190 m downhole depth, i.e., at c. 263–125 m a.s.l. (fig. 7). it is thus 138 m thick. a detailed log of the top zone is shown in fig. 15. the basal contact to the underlying dark mudstone is sharp, with only c. 10 cm basal breccia consisting of angular fragments in a brown matrix presumably of altered glass; a white carbonate-silicate mineral vein c. 1 cm in thickness cuts the basal zone (fig. 16j). the solid rock near the contact is very fine-grained but not strongly chilled. the main rock is relatively uniform in character throughout. it is a massive, fine-grained, grey rock with scattered orthopyroxene phenocrysts 2–5 mm long. the groundmass structure varies from nearly homogeneous over streaky to patchy with intermixed darker and lighter grey patches with bulbous and wispy shapes (fig. 16g). a detailed study of the mineralogy and chemistry of the body was conducted by lode et al. (2021). the native iron is finely disseminated through the rock and is not visible to the naked eye. sediment xenoliths up to 4 cm in size occur scattered through the length of the core; they consist of black mudstones and light yellowish sandstones with lobate and digested margins and reaction rims (fig. 16h–i). the lower c. 25 m of the core is broken up by long, vertical fractures. the top zone is c. 6 m thick (58.1–52 m) and is brecciated in a number of intervals with gradual transitions between them (fig. 15). the lowest zone (58.1–55.6 m) is greenish grey and more or less fragmental; it becomes less compact upwards, and the clasts become more easily visible. both clasts and matrix are finely vesicular (fig. 16f). at 55.6–55.4 m, the degree of oxidation increases upwards, and at 55.4–54.0 m, the rock is a striking, red-oxidised volcanic breccia with brick-red to reddish, dark grey to almost black, angular, vesicular clasts of highly variable size up to several centimetres, but most commonly around 1 cm. the matrix is whitish grey and carbonate-rich (fig. 16e). at 54.0–53.3 m, the volcanic breccia is grey to brownish grey, with a gradual transition to oxidised clasts in the upper 10 cm. at 53.3–52.8 m, the rock is a red-oxidised volcanic breccia of a character similar to that at the lower level (fig. 16d). at 52.8–52.0 m, the breccia is greenish grey and compact, with close-lying, irregular, mostly rounded, vesicular clasts up to more than 1 cm in size (fig. 16c). scattered orthopyroxene phenocrysts in the breccia clasts show that these consist of a rock similar to that of the massive lava. around 52 m, the rock has large clasts scattered in a clearly sedimentary matrix (fig. 16b). there is no sharp boundary between igneous rock and sediment, but we put the top of the igneous body at 52.0 m. 7 discussion 7.1 correlation of cores and the existence of a fault between them it is possible to correlate the major parts of the atane formation in cores fp94-4-5 and fp94-4-6 (fig. 17) and to extend the correlation to five short logs from the coastal cliffs north-west of qullissaaqqat (figs 2 and 3 and table 1). the correlation is primarily based on the sedimentary successions in both cores and is supported by the presence of coniacian dinoflagellate cysts in both fig. 15 detailed log of the brecciated top zone of the nunngarut igneous body and the overlying sediments. the positions of the photographs in fig. 16 are indicated. 50 m 51 52 53 54 55 56 57 58 • fig. 16a • fig. 16b • fig. 16c • fig. 16d • fig. 16e • fig. 16f https://doi.org/10.34194/geusb.v57.8361 http://www.geusbulletin.org/ larsen et al. 2024: geus bulletin 57. 8361. https://doi.org/10.34194/geusb.v57.8361 19 of 27 geusbulletin.org fig. 16 photographs of details of the nunngarut igneous body and immediately overlying sediments. beneath each piece, the downhole depth is indicated. positions in the core are shown in fig. 15. a: sediment (colluvium) with altered volcanic clasts in a weakly bedded matrix. b: large, altered volcanic clasts in a clearly sedimentary matrix, immediately above the top of the igneous body. c: grey volcanic breccia with close-lying, commonly vesiculated, volcanic fragments and little matrix. d: upper red breccia zone of more or less oxidised, volcanic fragments in a light, carbonate-rich matrix. e: lower red breccia zone; note vesiculation in the uppermost, large, angular fragment. f: dark grey, compact volcanic breccia with angular, finely vesiculated clasts. g: inhomogeneous, ‘wispy’ facies in the central part of the body. h: partly digested sandstone xenolith. i: partly digested mudstone xenolith. note the homogeneous facies of the host rock surrounding both xenoliths. j: the base of the igneous body. the rock is solid above the white vein, and the basal breccia is seen below the vein. c a b e d g i j f h 51.63 – 51.7 m 52.0 – 52.05 m 53.0 – 53.1 m 56.5 – 56.6 m 145.2 m 104.5 m 52.55 – 52.64 m 54.2 – 54.3 m 134.26 – 134.32 m 190.27 – 190.35 m https://doi.org/10.34194/geusb.v57.8361 http://www.geusbulletin.org/ larsen et al. 2024: geus bulletin 57. 8361. https://doi.org/10.34194/geusb.v57.8361 20 of 27 geusbulletin.org cores; the turonian–coniacian transition is present in core fp94-4-6 and probably present in core fp94-4-5 (figs 8, 10 and 17). correlation of similar successions of the qilakitsoq member has shown that the thick coarsening-upwards units are laterally continuous and, thus, are reliable markers for correlation (dam et al. 2009, fig. 46). the sedimentological correlation shows minor variations in thicknesses and facies, but these are known from previous studies of the formation and agree with studies of similar modern environments. all seven sections in fig. 17 include several coal beds. correlation to the coal layer exploited in the closed mine at the coast some 2 km farther north-west is uncertain due to poor exposures, but it is probably identical to layer b in section q-3. fig. 17 correlation diagram of the sedimentary facies associations in the two cores and five surface sections located in figs 2 and 3. 270 260 250 240 230 220 210 200 190 30 40 50 60 70 80 90 100 110 120 tr 130 140 143 28 50 40 35 tr 30 25 20 15 10 5 0 45 fp 94-4-5 fp 94-4-6 q-1 q-2 q-3 m-1 m-2 tr tr tr b co.co? co. a c legend marine dino�agellate cysts absent marine dino�agellate cysts present transgressive sandstone bed coal seams in coastal sections q and m coniacian above this level coniacian possibly above this level depositional environments and sedimentary structures: see legend to fig. 12 tr co? a, b, c co. https://doi.org/10.34194/geusb.v57.8361 http://www.geusbulletin.org/ larsen et al. 2024: geus bulletin 57. 8361. https://doi.org/10.34194/geusb.v57.8361 21 of 27 geusbulletin.org the correlation in fig. 17 conflicts with the present depths of the cores (fig. 7). the correlation indicates that the strata known from borehole fp94-4-6 and the coastal sections have been displaced approximately 90 m downwards relative to the strata in borehole fp944-5. the observed lack of the coniacian marker species chatangiella mcintyrei in the lower part of fp94-4-5 could perhaps be an effect of the very low density of dinoflagellate cysts in the samples, where a single specimen of chatangiella mcintyrei in a sample could make a difference (figs 8 and 10). on the balance of the evidence, in particular the mutual consistency of the two independent lines of evidence (facies correlation and palynology), we consider that the strata in core fp-94-4-6 have been displaced c. 90 m relative to the succession in core fp94-4-5 (fig. 7). the displacement process could have been faulting or landsliding. large parts of the coastal slopes in this region have repeatedly slipped, and, indeed, borehole fp94-4-5 was placed just outside the headscarp of a large rotational slide seen in fig. 2. this slide is relatively superficial, and the slide plane is located above the sediments in the coastal cliffs, so that the sediments below the slide plane are not deformed by the sliding (jackisch et al. 2022). however, the downward displacement of the strata in borehole fp94-4-6, and also in the coastal sections, has affected the strata down to at least 76 m below sea level (the bottom of hole fp94-4-6) and has left the sediments on both sides of the displacement plane near-horizontal and apparently undeformed. a very large interglacial landslide with a steep sliding plane could perhaps have caused the movement. a steep fault would also have the required effect. such a fault must strike roughly nw–se, which agrees with the nnw–sse to nw–se striking faults generated during the late rift phase in the early campanian to early maastrichtian (chalmers et al. 1999; dam et al. 2000; hopper et al. 2016, fig. 3.5; dam & sønderholm 2021, fig. 5). we consider faulting to be the most likely cause of the displacement. 7.2 atane formation the atane formation is largely unknown along the north-east coast of disko between asuk and nuugaarsuk, a stretch of c. 60 km (fig. 1), where exposures are scarce and have not been dated until now. the palynological results from the two cores show that the atane formation sediments between qullissaaqqat and qullissat are of late turonian–early coniacian age and belong to the qilakitsoq member (figs 4, 8 and 10). this result is unexpected because previous examinations of palynomorphs from the atane formation in outcrops from asuk to kussinerujuk have indicated a late albian to cenomanian age (bojesen-koefoed et al. 2007; pedersen & nøhr-hansen 2014) and thus the presence of the kingittoq member (fig. 4; dam et al. 2009). on nuussuaq, the areal distribution of the kingittoq and qilakitsoq members (fig. 5) appears to be relatively straightforward: the members occupy separate areas with the kingittoq member to the east and the qilakitsoq member to the west (probably overlying distal parts of the kingittoq member at depth). both members are exposed up to altitudes of 600–800 m, which indicates that a fault with downthrow to the west is located in the 4-km-wide, unexposed zone between kingittoq and paatuut (figs 1 and 5; a.k. pedersen et al. 2007a,b), as also indicated by, e.g., chalmers et al. (1999) and hopper et al. (2016, fig. 3.5). on the north coast of disko, the relation between the kingittoq and qilakitsoq members is probably governed by at least one and possibly more faults that are not exposed. in the well-exposed coastal cliff at asuk, the strata of the kingittoq member dip about 5° eastwards so that the general younging of the succession is towards south-east; a fault located south-east of qullissaaqqat, with a strike n–s or ne–sw and downthrow to the west, may explain why the coastal cliffs in the qullissat–qullissaaqqat area expose the younger qilakitsoq mb, whereas farther to the east, the older skansen member is exposed. several faults striking n–s to ne–sw are located in the vaigat strait, and some of these point directly towards qullissat and qullissaaqqat (chalmers et al. 1999; marcussen et al. 2002; hopper et al. 2016). these faults were most probably generated during the late rift phase. in any case, the faulting must be prevolcanic because the overlying volcanic rocks are generally unfaulted. a better constrained structural analysis of the atane formation on disko awaits more age determinations of the undated parts (fig. 5). 7.3. eqalulik formation the boundary between the cretaceous atane formation and the overlying danian eqalulik formation is located at 205 m in core fp94-4-5 (fig. 12) and represents a hiatus of c. 24 million years (cf. above). the cross-bedded sandstone below 205 m is interpreted as part of the qilakitsoq member of the atane formation. the medium-grained to fine-grained, structureless sandstone above is interpreted as deposited above a transgressive surface of erosion, formed during the rapid transgression at the base of the eqalulik formation (tss8), discussed by dam et al. (1998b). this transgression is only recorded in some offshore wells (nøhr-hansen et al. 2002; fensome et al. 2016). only a single small exposure of the eqalulik formation is known on disko. in the coastal cliff in the landslide at asuk, 40 cm of tuffaceous marine mudstone https://doi.org/10.34194/geusb.v57.8361 http://www.geusbulletin.org/ larsen et al. 2024: geus bulletin 57. 8361. https://doi.org/10.34194/geusb.v57.8361 22 of 27 geusbulletin.org with dinoflagellate cysts overlies the top of a submarine lava flow and is covered by a subaerial flow, both of the asuk member (piasecki et al. 1992; pedersen et al. 2017, fig. 95). the mudstone is a late part of the eqalulik formation, whereas the main part must be situated below sea level. a palaeogeographic reconstruction of the marine embayment in which the formation was deposited (pedersen et al. 2017, fig. 136) indicates that it should be present in the qullissaaqqat area, which is now confirmed. the lack of volcanic material in the core may reflect the long distance to the major volcanic eruption sites in the north-west (fig. 6) and the general low explosivity of the vaigat formation. the nearest other locality with a complete section through the eqalulik formation is exposed at ataata kuua on nuussuaq, 30 km due north of qullissaaqqat. 7.4 the nunngarut lava flow the strongly red-oxidised top breccia of the body shows that the magma has been exposed to open air; it is, thus, not a sill but a lava flow. the two grey-and-red intervals in the top breccia show that the semi-consolidated top crust has been broken up in slabs that were stacked during the final flowage of the magma. very similar lava flows with stacked and tumbled slabs of crust on the top are described from the deccan province by kale et al. (2022). at the base of the flow, the sediment has been strongly heated, as shown by the high vitrinite reflectance (table 2). the long, vertical fractures in the lower c. 25 m of the core are interpreted as columnar jointing, indicating that the lower part of the flow interacted with water. the conclusion that the mineralised body is not a sill but a lava flow reduces its economic potential according to the noril’sk mineralisation model (lightfoot & hawkesworth 1997) because the magma throughflow during lava emplacement is expected to be limited. however, its thickness, distribution and volume will still be of interest for future exploration, and the nuungarut flow is significantly thicker than any other flow or sill known from disko and nuussuaq. similar iron-bearing magnesian basaltic-andesite flows from the asuk member have thicknesses up to 45 m at asuk and 60 m in the kuugannguaq valley (pedersen 1985, fig. 9; pedersen et al. 2017, figs 94 and 98). a dacite flow from the nordfjord member of the maligât formation is 120 m thick (pedersen et al. 2018, fig. 104c). the thickest flows of basaltic composition on disko and nuussuaq are 60–80 m thick; these are commonly flows that have run into water-filled depressions and become ponded there. as the nuungarut flow is only exposed in a coastal section of c. 4.7 km length, its extent (present as well as original) cannot be known. the flow most probably extends to the west beneath the high mountain ridge where it cannot be traced (fig. 3; jackisch et al. 2022). on the other side of the mountain ridge, flows of the asuk member are exposed in both sides of the kuugannguaq valley. in this valley, the main area with asuk member reaches its south-easternmost limit 20 km due west of qullissaaqqat (fig. 6). from southern kuugannguaq, the south-east limit of the member runs north-east beneath the mountain ridge and reappears in the landslipped areas south of asuk, 12 km north-west of the nunngarut flow at qullissat (fig. 6). in kuugannguaq and at asuk, there are ironand graphite-bearing flows with compositions very similar to that of the nunngarut flow, in particular ‘composite lava 1’ in kuugannguaq (table 3). no direct eruption sites in the form of craters or feeders for the asuk member flows are known, but the main production area must be situated centrally within the area of distribution of the asuk member on disko (fig. 6). if the nunngarut flow was erupted from the central area, it must have travelled at least 15 km and probably closer to 20–25 km to reach qullissaaqqat. alternatively, the nunngarut flow was produced from a separate eruption site about 20 km east of the main production sites, in an area where no magmas were previously produced; however, this is not likely to lead to so closely similar magma compositions as shown in table 3. we consider it more likely that the magma was erupted from the main production area and flowed the same distance on the ground, aided by gravity. long travel distances for such relatively siliceous lava flows are known from the nordfjord member of the maligât formation in western disko. here, the mellemfjord lava flow is a composite, iron-bearing marker flow with basaltic to andesitic composition that covers an area of 28 by 15 km, i.e. more than 400 km2. it has a thickness up to 88 m and a volume of more than 14 km3 and is one of the largest flows in the nuussuaq basin (pedersen 1977; pedersen & ulff-møller 1987; pedersen et al. 2018, pp. 147–151). the known volume of the nunngarut flow is approximately 5 km (length) × 2 km (width) × 0.13 km (thickness) = 1.3 km3. however, including its probable extension beneath the younger volcanics and its long travel distance, it must have a considerably larger volume. moreover, the close compositional similarity between the nunngarut flow and the iron-bearing flow in kuugannguaq suggests the possibility that the two flows were produced during the same eruptional event, and in this case, the erupted magma volume may be the largest in the asuk member. 7.4.1 emplacement of the nunngarut lava flow at the time of eruption of the asuk member, the lava plateau formed an emergent land area in north-western disko and western nuussuaq. in central disko, https://doi.org/10.34194/geusb.v57.8361 http://www.geusbulletin.org/ larsen et al. 2024: geus bulletin 57. 8361. https://doi.org/10.34194/geusb.v57.8361 23 of 27 geusbulletin.org the volcanic rocks were banked up against the emergent disko gneiss ridge. to the east and south-east, the volcanic rocks prograded into a narrowing marine embayment connected to the main sea north of nuussuaq (pedersen et al. 2017, fig. 136). the eastern shore of this embayment received mud and sand from the paleocene successor of the large cretaceous river and delta system, and at the sea floor, the synvolcanic mudstones of the eqalulik formation accumulated. the surface of the prevolcanic sediments was irregular. when the vaigat formation was erupted, highs with palaeovalleys in the old sediments were locally emergent, for example, at naajannguit 35 km north-west of qullissat, where such a valley was filled with subaerial lava flows of the naujánguit member of the vaigat formation, some of the first lava flows that reached disko (pedersen et al. 2002). we envisage that the nunngarut lava flow was erupted in the main magma production area between kuugannguaq and the north-east coast of disko (fig. 6). in this area, accumulated eruption products could have formed a relatively elevated land area, perhaps even one or more local volcanic edifices. lavas that flowed to the west and south would cover older subaerial lava flows and eventually reach the disko gneiss ridge, as recorded by pedersen et al. (2005). lavas that flowed to the east and north-east would enter the sea and become brecciated like the first flow at asuk (pedersen et al. 2017, fig. 94). lavas that flowed into shallow water and whose top stayed above the water level would maintain coherence but develop colonnades and entablatures like the second flow at asuk and the nunngarut flow. a low in the east-sloping sediment surface could have guided the nunngarut flow for perhaps as much as 20 km to the south-east and into the sea. a coast-near local basin could provide space for ponding of the flow to create its extraordinary thickness. the flow was not brecciated but attained columnar jointing, and its oxidised surface stayed well above the sea level, at least at the drill site. the ability of the lava flow to reach so far and attain such a thickness, even by ponding, requires that the erupted volume was considerable. 7.5 atanikerluk formation the c. 1 m of grey to violet-grey, weakly bedded volcaniclastic sediment with siliciclastic grains overlying the top of the nunngarut lava flow is interpreted as an outwash/ colluvial deposit of altered volcanic material and disintegrated quartzofeldspathic sediment washed onto the weathered lava surface from surrounding exposures. the deposit is considered to be of local extent. the overlying c. 20 m of cross-bedded sandstone and subordinate mudstone with thin horizons of organic matter of fluvial and lacustrine origin probably had a more extensive distribution. the sand-dominated lithologies in the fp94-4-5 core, with sparse reworked cretaceous dinoflagellate cysts and paleocene spores and pollen, are comparable to those of the akunneq member of the lower atanikerluk formation, which is exposed in eastern disko between nuugaarsuk and pingu, 36–50 km south-east of qullissaaqqat. here, the akunneq member has thicknesses of 145–165 m, decreasing westwards, and comprises fluvial sandstones with thin interbedded mudstones (dam et al. 2009, figs 128, 132, 133). the lower part contains reworked cenomanian pollen and few paleocene spores and pollen, whereas the upper part contains more paleocene spores and pollen. we tentatively refer the sediments above the nunngarut lava flow in core fp94-4-5 to the akunneq member; the thickness of 20 m is a result of erosion. the akunneq member in eastern disko is not in contact with the volcanic rocks; however, the member is interpreted to correlate with the lower part of the naujât member on nuussuaq, which, in turn, is coeval with the ordlingassoq member of the vaigat formation (dam et al. 2009, p. 149), i.e. close in time to but slightly younger than the nunngarut lava flow. the sediments overlying the nunngarut lava flow in fp94-4-5 present a rare glimpse of the atanikerluk table 3 chemical compositions of the nunngarut lava flow and a lava flow from the asuk member in kuugannguaq. nunngarut flow1 kuugannguaq2 depth in core ggu no. 89–90 m 176736 major elements, wt% sio2 54.94 55.01 tio2 1.14 1.11 al2o3 15.30 15.19 feo* 10.18 9.87 mno 0.14 0.13 mgo 7.26 7.42 cao 8.24 8.20 na2o 1.73 2.14 k2o 0.93 0.80 p2o5 0.14 0.11 100.00 100.00 s wt% 1.04 0.91 trace elements, ppm cr 547 627 ni 433 435 co 81 57 cu 232 329 zn 19 47 rb 27 26 ba 218 231 sr 186 195 nb 7 7 zr 131 136 y 21 24 major elements were calculated volatile-free. 1analysis from olshefsky et al. (1995). 2composite lava 1, 70°6.5’n, 53°41.1’e, 357 m a.s.l. 2analysis from pedersen et al. (2017). 2photograph in pedersen et al. (2017), fig 98. https://doi.org/10.34194/geusb.v57.8361 http://www.geusbulletin.org/ larsen et al. 2024: geus bulletin 57. 8361. https://doi.org/10.34194/geusb.v57.8361 24 of 27 geusbulletin.org formation in north-eastern disko, where exposures are very rare. the only other nearby in situ exposure of the atanikerluk formation sediments occurs 4.5 km northwest of fp94-4-5 at c. 700 m altitude (fig. 1b). these are mudstones tentatively referred to the assoq member (upper atanikerluk formation) by a.k. pedersen et al. (2018, p. 63) and are time-equivalent to the lower rinks dal member of the maligât formation. a complete section through the atanikerluk formation in disko west of nuugaarsuk is still not known. 8 conclusions the cretaceous sedimentary strata between qullissat and qullissaaqqat comprise a succession of delta deposits more than 130 m thick. the dinoflagellate cyst assemblages indicate a late turonian to early coniacian age, and the succession is referred to the qilakitsoq member of the atane formation. this member has not been recorded on disko before. the cretaceous sediment successions in the two drill cores at qullissaaqqat are mutually correlatable and are further correlatable with five short sections in the coastal cliffs between qullissaaqqat and qullissat. thereby, the coal quarried at ritenbenk’s coal quarry and mined at qullissat is dated to be of early coniacian age. the atane formation (qilakitsoq member) at qullissaaqqat is younger than the atane formation in both eastern disko (skansen member) and north-western disko (kingittoq member). this structure indicates the presence of one or more hidden faults, including one east of qullissaaqqat. more precise location of the fault, or faults, requires better knowledge of the ages of the formation over long distances of the north-east coast of disko. the range charts (figs 8 and 10) show that marine dinoflagellate cysts are present in roughly half of the samples from the atane formation. this indicates that the cores represent a part of the delta where non-marine and marine depositional environments changed repeatedly. palynological samples from coal seams b and c from outcrop section q-3 are all dominated by spores from ferns and bryophytes. this suggests that both coal seams represent relatively open mires. the pollen in the samples may come from trees and bushes on higher elevated areas or the hinterland. available vitrinite reflectance data from the area consistently suggest a maximum depth of burial of less than 2 km when allowing for a slightly higher than normal geothermal gradient related to the igneous activity in the region (bojesen-koefoed et al. 1997, 2001; pedersen et al. 2006). the danian eqalulik formation in core fp94-4-5 consists of 4 m of sand and 11 m of mudstone, of which at least the upper 10 cm is strongly baked by the overlying lava flow. the presence of the formation is in accordance with the palaeogeographic reconstruction of the marine embayment in which the formation was deposited (pedersen et al. 2017, fig. 136). this is the only section through the eqalulik formation known from disko. the native-iron-bearing nunngarut igneous body is a lava flow belonging to the asuk member of the vaigat formation. the change of status from ‘assumed sill’ to subaerial lava flow diminishes its economic potential because the magma throughflow in a lava flow is expected to be limited. however, the thickness, distribution and volume of the flow will still be of interest for future exploration. it is the thickest flow known from west greenland and was most probably erupted in a central eruption area for the asuk member between the kuugannguaq valley and the present north-east coast of disko. it has run towards south-east over sloping terrain, possibly guided through a surface low, for a distance of perhaps up to 20 km to the sea where it ponded and, at least in the area it was drilled, completely displaced the water so that the red-oxidised top was maintained. the erupted volume must have been much larger than the presently known volume of only 1.3 km3. the atanikerluk formation in core fp94-4-5 consists of c. 1 m of locally derived colluvial material deposited on the weathered lava surface, and at least 21 m of fluvial to lacustrine, sand-dominated deposits referred to the akunneq member, which is time-equivalent to the ordlingassoq member of the vaigat formation. the nearest comparable exposures occur 36 km south-east of qullissaaqqat, illustrating the extreme scarcity of sediment exposures on the north-east coast of disko. acknowledgements technical assistance by geus personnel jens gregersen, annette ryge, charlotte olsen, carsten guvad, stine øckenholt and mads porse is highly appreciated. kristian svennevig is thanked for the photograph in fig. 2. we thank henrik tirsgaard and james b. riding for helpful and constructive reviews. additional information funding statement all expenses in connection with this work were covered by the geological survey of denmark and greenland (geus). author contributions lml, gkp, akp: conceptualisation, investigation (geology), writing (lead), editing. hnh, sl: investigation (palynology), writing, editing. jbk: investigation (vitrinite reflectance, organic geochemistry), writing, editing. evs: investigation (photogrammetry), writing, editing. competing interests the authors declare no competing interests. additional files none. https://doi.org/10.34194/geusb.v57.8361 http://www.geusbulletin.org/ larsen et al. 2024: geus bulletin 57. 8361. https://doi.org/10.34194/geusb.v57.8361 25 of 27 geusbulletin.org references andrews, s., vosgerau, h. & bojesen-koefoed, j. 2022: the sedimentology and depositional environments of the bastians dal and muslingebjerg formations: evidence for the earliest phases of jurassic rifting in north-east greenland. geus bulletin 49, 8311. https://doi. org/10.34194/geusb.v49.8311 bailey, d. & biostrat stratigraphic consultancy 2021: late cretaceous zonation. http://www.biostrat.org.uk/index.html# barker, c.e. & pawlewicz, m.j. 1994: calculation of vitrinite reflectance from thermal histories and peak temperature: a comparison of methods. in: mukhopadhyay, p.k.a.d. & dow, w.g. 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(eds): geology of greenland, 507-516. copenhagen: geological survey of greenland. https://doi.org/10.22008/gpub/38218 shekhar, s.c., frandsen, n. & thomsen, e. 1982: coal on nûgssuaq, west greenland. unpublished report, geological survey of greenland, 82 pp + appendices. steenstrup, k.j.v. 1874: om de kulförende dannelser på öen disko, hareöen og syd-siden af nûgssuak’s halvöen i nord-grönland. videnskabelige meddelelser fra den naturhistoriske forening i kjöbenhavn for aaret 1874, 75–112 + 3 plates. steenstrup, k.j.v. 1875: om de nordenskiöldske jærnmasser og om forekomsten af gedigent jærn i basalt. videnskabelige meddelelser fra den naturhistoriske forening i kjöbenhavn for aaret 1875, 284–306 + 2 plates. steenstrup, k.j.v. 1877: on the non-meteoric origin of the masses of metallic iron in the basalt of disko in greenland. mineralogical magazine 1, 143–148. https://doi.org/10.1180/minmag.1877.005.1.02 steenstrup, k.j.v. 1882: om forekomsten af nikkeljern med widmannstättenske figurer i basalten i nord-grønland. meddelelser om grønland 4(3), 113–131. steenstrup, k.j.v. 1900: beretning om en undersøgelsesrejse til øen disko i sommeren 1898: meddelelser om grønland 24, 249–306 + plates. storey, m., duncan, r.a., pedersen, a.k., larsen, l.m. & larsen, h.c. 1998: 40ar/39ar geochronology of the west greenland tertiary volcanic province. earth and planetary science letters 160, 569–586. https:// doi.org/10.1016/s0012-821x(98)00112-5 styrelsen for dataforsyning og infrastruktur 2023: højdemodel grønland. https://dataforsyningen.dk/data/4780 svennevig, k., owen, m.j., citterio, m., nielsen, t., rosing, s., harff, j., endier, r., morlighem, m. & rignot, e. 2023: holocene gigascale rock avalanches in vaigat strait, west greenland – implications for geohazard. geology. https//doi.org/10.1130.g51234.1 talling, p.j., masson, d.g., sumner, e.j. & malgesini, g. 2012: subaqueous sediment density flows: depositional processes and deposit types. sedimentology 59, 1937–2003. https://doi.org/10.1111/j.1365-3091. 2012.01353.x https://doi.org/10.34194/geusb.v57.8361 http://www.geusbulletin.org/ https://doi.org/10.1306/212f7b44-2b24-11d7-8648000102c1865d https://doi.org/10.1306/212f7b44-2b24-11d7-8648000102c1865d https://doi.org/10.22008/gpub/38218 https://doi.org/10.1180/minmag.1877.005.1.02 https://doi.org/10.1016/s0012-821x(98)00112-5 https://doi.org/10.1016/s0012-821x(98)00112-5 https://dataforsyningen.dk/data/4780 http://doi.org/10.1130.g51234.1 https://doi.org/10.1111/j.1365-3091.2012.01353.x https://doi.org/10.1111/j.1365-3091.2012.01353.x a hidden succession revealed: cretaceous and paleocene sediments and a native-iron-bearing lava flow in cores near qullissat, disko, west greenland 1 introduction 2 history 3 geological setting 4 geology of the formations represented in the drill cores 4.1 sedimentary rocks 4.1.1 atane formation 4.1.2 eqalulik formation 4.1.3 atanikerluk formation 4.2 igneous rocks 4.2.1 vaigat formation 5 material and methods 6 results 6.1 palynology and ages 6.1.1 core fp94-4-5 6.1.2 core fp94-4-6 6.1.3 palynology of coal seams b-c in the coastal cliff at qullissaaqqat 6.2 vitrinite reflectance and organic geochemistry 6.3 atane formation 6.3.1 fluvial and distributary channel sandstones 6.3.2 delta plain deposits 6.3.3 transgressive shoreface sandstones 6.3.4 delta front deposits 6.4 eqalulik formation 6.5 atanikerluk formation 6.6 the nunngarut igneous body 7 discussion 7.1 correlation of cores and the existence of a fault between them 7.2 atane formation 7.3. eqalulik formation 7.4 the nunngarut lava flow 7.4.1 emplacement of the nunngarut lava flow 7.5 atanikerluk formation 8 conclusions acknowledgements additional information references figures fig. 1 geological maps of parts of the nuussuaq basin. a: overview map of disko and nuussuaq, with names mentioned in the text and the position of fig. 1b indicated. b: map of the area around qullissat and qullissaaqqat, with boreholes fp94-4-5 and fp94-4-6 indicated. two exposures of the iron-bearing igneous body studied here are labelled ‘fe’. other labels are part of the geological divisions on the map. excerpt from the geological map 1:100 000 qutdligssat (old spelling, rosenkrantz et al. 1976), reproduced to scale, i.e. 1 cm = 1 km. fig. 2 oblique aerial photograph of the coastal area of disko between the closed coal mine (at right) and qullissaaqqat (immediately left of the picture). ritenbenk’s old coal quarry is clearly seen from the air; seen on the ground it is inconspicuous. the positions of boreholes fp94-4-5 and fp94-4-6 and coastal sections q-1, q-2, q3, m-1 and m-2 are indicated. exposures of the nunngarut igneous body (fe) in situ are annotated with white lines to the right and left; a third exposure is seen in a landslide at the coast at nunngarut. the height of the coastal cliffs is around 50 m, and the height of the plateau fig. 3 photogrammetrically measured and interpreted section along the coast of disko between qullissat (at right) and qullissaaqqat (immediately left of the section). exposures of the iron-bearing nunngarut igneous body are red. the vertical extents of boreholes fp94-4-5 and fp94-4-6 are indicated. compare fig. 2, but note that the perspectives are very different because the section possesses no single point of view. modified from fig. 48 in pedersen fig. 4 stratigraphy and age of the sedimentary and volcanic formations of the nuussuaq basin on disko and nuussuaq. on the age scale to the left, rings around two numbers and change of colours indicate changes in the scale. tss-1 to tss-8 are tectonostratigrahic sequences (dam & sønderholm 2021). ku.: kussinerujuk member. members of the volcanic formations are a: anaanaa member; n: naujánguit member; o: ordlingassoq fig. 5 map showing schematically the known extents of the four members of the atane formation of the nuussuaq group. the lines do not indicate member boundaries but knowledge limitations: members can only be assigned in the areas with numbers 1 to 4. the atane formation is not present in western and northern nuussuaq, except in the area labelled 1 and 3. fig. 6 map showing the distribution of the individual units of the asuk member. ce: inferred central eruption area for the asuk member on disko. the large, light grey circular area is the extent of the graphite andesite tuff layer erupted from the ilugissoq volcano. the green areas show exposures of the vaigat formation, and the green lines are the eastern and south-eastern depositional boundaries of the large naujánguit fig. 7 simplified overview logs of the investigated cores fp94-4-5 and fp94-4-6 showing their position relative to sea level (scale at left). downhole depths are shown along each core. correlation of the cretaceous sediments in the two cores is based on their sedimentological logs (see below). a characteristic correlatable drowning surface is shown in blue. the surface is situated c. 60 m above sea level in core fp94-4-5 and c. 30 m below sea level in core fp94-4-6, indicating a vertical displacement between the drill sites of c. 90 m along a fault. the fault is schematically indicated on the drawing; its precise position and orientation are uncertain. fig. 8 range chart of dinoflagellate cysts, spores and pollen (sp), acritarchs, fungi (fu) and freshwater algae (al) in core fp94-4-5. biostratigraphically important palyno-events are shown to the right. fo: first occurrence. lo: last occurrence. fig. 9 palynomorphs from core fp94-4-5. same magnification for all figures; scale bar in aa = 20 μm. each figure represents species name, sample no., slide no., england finder coordinate and photo no. a: taurodinium granulatum, sample 393016-100, slide 2, w52-3, photo 2709. b: taurodinium granulatum, sample 393016-100, slide 3, p28-4, photo 2711. c: taurodinium granulatum, sample 393016-100, slide 2, w54-1, photo 2713. d: pediastrum sp., sample 393016-100, slide 2, f41-1, photo 2710. e: taurodinium granulatum, sample 393016-122, slide 2, h57-1, photo 2708. f: dinocyst sp. 3 nøhr-hansen fig. 10 range chart of dinoflagellate cysts, spores and pollen, and acritarchs (ac) in core fp94-4-6. biostratigraphically important palyno-events are shown to the right. fo: first occurrence. lo: last occurrence fig. 11 palynomorphs from core fp94-4-6. same magnification for all figures; scale bar in t = 20 μm. each figure represents species name, sample no., slide no., england finder coordinate and photo no. a: heterosphaeridium difficile, sample 393017-15, slide 5, e50-3, multifocus. b: chatangiella mcintyrei, sample 393017-15, slide 4, g48-4, photo 2741. c: raphidodinium fucatum, sample 393017-12, slide 4, w34-4, photo 2740. d: chatangiella mcintyrei, sample fig. 12 sedimentological logs of cores fp94-4-5 and fp94-4-6. fig. 13 photographs of the atane formation sediments in the cores. core diameter is 36 mm. a: fineto medium-grained, cross-bedded sandstone, locally with coalified plant debris, interpreted as fluvial or distributary channel facies. core fp94-4-5, depth c. 237.70–237.50 m. b: delta plain facies, dark grey mudstone with lenses or thin layers of coal. core fp94-4-5, depth c. 238.70–238.55 m. c: delta plain mudstone erosively overlain by bioturbated fig. 14 photographs of the eqalulik and atanikerluk formation sediments in the cores. a–d: eqalulik formation; e–g: atanikerluk formation; all from core fp94-4-5, diameters 36 mm. a: structureless sandstone with very small mudstone clasts in a mudstone matrix. lower part of the eqalulik formation, depth c. 201.60–201.45 m. b: dark grey mudstone with sharp-based streaks and thin layers of pale, very fine-grained sand and coarse silt. the thickest layers (3–7 mm) are faintly graded. red-brown colour is due to cementation. depth c. 200.55–200.40 m. c: mudstone facies characterised by thin layers of angular, light grey grains in a mudstone matrix, interpreted as deposited from thin mudflows with sufficient matrix strength to transport the sandto granule-sized grains. depth c. 198.95–198.80 m. note the sharp boundary to homogeneous mudstone at c. 198.85 m. d: grey, silt-streaked mudstone, marine deep-water deposit, upper part of the eqalulik formation, depth c. 191.95–191.80 m. e: cross-bedded, grey sandstone with tiny mudstone clasts. atanikerluk formation, depth c. 48.08–47.95 m. f: cross-bedded, dark grey sandstone with numerous tiny mudstone clasts, depth c. 46.45–46.30 m. g: angular and rounded, moderately sorted clasts in a matrix od sand and mud, interpreted as colluvial deposits. depth c. 36.90–36.75 m. fig. 15 detailed log of the brecciated top zone of the nunngarut igneous body and the overlying sediments. the positions of the photographs in fig. 16 are indicated. fig. 16 photographs of details of the nunngarut igneous body and immediately overlying sediments. beneath each piece, the downhole depth is indicated. positions in the core are shown in fig. 15. a: sediment (colluvium) with altered volcanic clasts in a weakly bedded matrix. b: large, altered volcanic clasts in a clearly sedimentary matrix, immediately above the top of the igneous body. c: grey volcanic breccia with close-lying, commonly vesiculated, volcanic fragments and little matrix. d: upper red breccia zone of more or less oxidised, volcanic fragments in a light, carbonate-rich matrix. e: lower red breccia zone; note vesiculation in the uppermost, large, angular fragment. f: dark grey, compact volcanic breccia with angular, finely vesiculated clasts. g: inhomogeneous, ‘wispy’ facies in the central part of the body. h: partly digested sandstone xenolith. i: partly digested mudstone fig. 17 correlation diagram of the sedimentary facies associations in the two cores and five surface sections located in figs 2 and 3. tables table 1 data for investigated drill cores and surface sections in the qullissat area. table 2 vitrinite reflectance (ro) and screening data for organic-rich samples below and above the nunngarut lava flow. table 3 chemical compositions of the nunngarut lava flow and a lava flow from the asuk member in kuugannguaq. research article dolan et al. 2022: geus bulletin 52. 8325. https://doi.org/10.34194/geusb.v52.8325 1 of 16 delivering seabed geodiversity information through multidisciplinary mapping initiatives: experiences from norway margaret f.j. dolan*1 , reidulv bøe1 , lilja r. bjarnadóttir1 1geological survey of norway (ngu), trondheim, norway abstract geology is a core component of two major multidisciplinary seabed-mapping initiatives in norway (mareano, marine base maps for the coastal zone). helped by norway’s nature diversity act, which acknowledges geological and landscape diversity alongside biodiversity, geological information has gained recognition nationally as part of an essential foundation for knowledge-based management, both in the coastal zone and offshore. recently, international focus on the united nations sustainable development goals has led to the proposal of essential geodiversity variables, a framework for geological (geodiversity) information, intended to stand alongside essential variables already defined for climate, biodiversity and oceans (limited to ocean physics, biochemistry, biology and ecosystems). here, we examine to what extent map products from the geological survey of norway generated under these multidisciplinary mapping initiatives fit within this framework of essential geodiversity variables, and how well it is suited to information on marine geodiversity. although we conclude that the framework is generally a good fit for the marine-relevant essential geodiversity variable classes (geology and geomorphology), we examine opportunities for further highlighting quantitative geodiversity information. we present preliminary examples of substrate diversity and morphological diversity and discuss our experience of geological mapping as part of multidisciplinary initiatives. we highlight many benefits, which far outweigh any perceived or real compromises of this approach in monetary, practical and scientific terms. *correspondence: margaret.dolan@ngu.no received: 23 jun 2022 accepted: 02 dec 2022 published: 29 dec 2022 keywords: geodiversity, geomorphons, marine geology, seabed mapping, sediment grain size abbreviations: cmecs: united states coastal and marine ecological classification standard egvs: essential geodiversity variables eovs: essential ocean variables evs: essential variables gdcs: geodiversity components imr: institute of marine research mareano: marine areal database for norwegian sea areas ngu: geological survey of norway nin: nature in norway sdgs: sustainable development goals geus bulletin is an open access, peerreviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: kaskela anu (geological survey of finland), verner brandbyge ernstsen (geological survey of denmark and greenland) reviewed by: joseph j bailey (anglia ruskin university, uk), lars erikstad (norwegian institute for nature research, norway) funding: see page 14 competing interests: see page 14 additional files: see page 14 1. introduction geodiversity (gray 2004) is generally regarded as the abiotic equivalent of biodiversity. encompassing diversity in rocks, sediments, landforms and physical processes that underpin our environment, or more simply the diversity of geological and geomorphological phenomena in a defined area (johansson et al. 2001), it is equally applicable in the marine realm as on land. despite the term ‘geodiversity’ being coined almost 30 years ago (sharples 1993; wiedenbein 1993), the importance of geodiversity remains far less well acknowledged or celebrated (gray 2021) than its biotic counterpart ‘biodiversity’, which gained far more political traction and public interest. milton’s (2002) statement ‘diversity in nature is usually taken to mean diversity of living nature …’, which has been highlighted by several authors on geodiversity (e.g. gray 2011; brilha et al. 2018), remains just as true today, not least in relation to marine benthic habitats. areas of rich plant and animal life on the seabed are frequently highlighted as biodiversity hotspots attracting the interest of nature conservation, whilst abiotic diversity, and specifically geodiversity, occurring over spatial scales beyond the camera lens, attracts less interest. geodiversity and biodiversity are often linked, with diversity hotspots often co-located, but this is not always the case. over nearly two decades of multidisciplinary seabed mapping in norway, we have frequently observed, particularly from underwater video surveys, that changes in seabed geology coincide with changes in the associated benthic communities. however, we https://doi.org/10.34194/geusb.v52.8325 https://orcid.org/0000-0003-4405-9277 https://orcid.org/0000-0002-9486-3530 https://orcid.org/0000-0002-3152-8348 mailto:margaret.dolan@ngu.no dolan et al. 2022: geus bulletin 52. 8325. https://doi.org/10.34194/geusb.v52.8325 2 of 16 www.geusbul let in.org also observe geological diversity that is important to highlight regardless of any known connection with biodiversity, as well as in areas where the seabed biota is undocumented. even though a connection between geodiversity and biodiversity is generally recognised, it is not so easily quantified, at least in consistent terms. this is partly linked to inconsistencies in the use of terminology and to what extent users perceive a need to align with concepts of biodiversity – see recent summary by gray (2021) and insight into potential oversimplification issues from erikstad (2013). regardless of any ongoing terminology debates, there seems to be growing interest in using geodiversity as a surrogate for biodiversity (e.g. hjort et al. 2012; tukiainen et al. 2016), whilst potential cost savings and added value of such an approach have been noted in terrestrial settings (bailey et al. 2017). this is a particularly attractive prospect in the marine realm, where accessibility and high associated costs with offshore surveys often limit the taxonomic inventories required to quantify biodiversity. by contrast, at least certain components of geodiversity can be mapped with the aid of remote sensing, preferably supplemented with limited in-situ observations (ground-truthing). in this sense, geodiversity mapping may offer an important foundation for identifying areas of conservation priority. this adds weight to the argument that geodiversity should be a priority for nature conservation (e.g. chakraborty & gray 2020). geodiversity has recently attracted attention in relation to the 17 united nations sustainable development goals (sdgs; e.g. brilha et al. (2018)) with the development of the essential geodiversity variables (egvs) framework proposed by schrodt et al. (2019) to supplement earlier essential variables (evs) defined for climate, biodiversity and oceans (e.g. bojinski et al. 2014). in relation to seabed mapping, we stress that the essential ocean variables (eovs) currently defined are based on the global ocean observing system (2022), as such they do not include variables relevant to seabed geology but are limited to ocean physics, biochemistry, biology and ecosystems. furthermore, they include only two variables of possible interest for seabed mapping in high latitudes: seagrass cover and composition as well as macroalgal canopy cover and composition, with coral reefs currently focussed on tropical rather than cold-water corals. egvs are defined by schrodt et al. (2019) as abiotic state and process variables that relate to geology, geomorphology, soils and hydrology and which are: 1. relevant to natural resource management and human well-being, conservation or ecology. 2. complementary to the other suites of evs. 3. feasible and cost effective to measure. whilst soils and hydrology are less relevant in the marine environment, we can still use the remainder of the proposed egv framework and assess its application in settings other than the terrestrial applications for which it was first conceptualised. the egv concept is still relatively new, and we may expect some refinement and further development over the coming years, perhaps including adaptation towards seabed mapping. establishing egvs as part of a suite of evs promotes the need to consider geodiversity as a core component of nature, alongside biodiversity, and to do so through consistent terminology. without incorporating geodiversity, we risk undervaluing nature (gray 2012) and may fail to recognise important geosystem services derived from geodiversity (gray 2021). here, we aim to assess to what extent the major seabed-mapping initiatives in norway currently deliver geodiversity information in relation to the egv framework. we restrict our focus to map products published and/or developed by the geological survey of norway (ngu). specifically, we aim to: • assess to what extent existing ngu marine geology products from norway’s multidisciplinary seabed-mapping initiatives fit into the egv framework. • provide some examples of how selected ngu products can better highlight geodiversity. • discuss the extent to which mapping geology as part of multidisciplinary seabed-mapping initiatives helps delivery of geodiversity information and the relevance of the egv framework to this. 2. geodiversity as part of multidisciplinary seabed-mapping initiatives in norway norway has made great strides in seabed mapping over the past couple of decades, benefiting from improvements in survey technology and it infrastructure over the same period, which have been so fundamental in supporting the acquisition of increased knowledge of the seabed. government and local or regional funding supports for these mapping initiatives have been substantial, reflecting the importance of coastal and offshore resources to the norwegian economy. here, we outline two of the largest initiatives currently underway, both of which ngu is a core partner. 2.1 mareano the norwegian seabed mapping programme mareano (marine areal database for norwegian sea areas) started in 2005 with a focus on offshore mapping. the programme is government-funded and receives annual contributions from two ministries (the ministry of https://doi.org/10.34194/geusb.v52.8325 http://www.geusbulletin.org dolan et al. 2022: geus bulletin 52. 8325. https://doi.org/10.34194/geusb.v52.8325 3 of 16 www.geusbul let in.org trade, industry and fisheries and the ministry of climate and environment) through the national budget. since 2005, nearly 1.4 billion nok (c. 135 million eur) have been invested in mareano’s marine mapping. mareano maps bathymetry, seabed substrates, biodiversity, habitats and pollution in seabed sediments. the multidisciplinary mapping is carried out by three collaborating institutions: the norwegian hydrographic service (part of the norwegian mapping authority), ngu and the institute of marine research (imr). since 2005, the seabed has been surveyed using acoustic remote sensing (multibeam bathymetry, backscatter, water-column data and sub-bottom-profiler data), whilst geology, biology and chemistry have been mapped via in situ video surveys and physical sampling. as an example, seabed sediments (grain size) maps have been made for areas covering 270 000 km2. until 2010, the programme focussed its efforts on the barents sea, before moving to include the norwegian sea and areas around svalbard. in 2019, mareano mapped about 69  000 km2 in the deeper parts of the norwegian sea, including several areas on the mid-atlantic ridge using acoustic remote sensing. follow-up in situ surveys are planned in the coming years to ground truth the acoustic data and acquire more detailed visual and acoustic data using underwater survey platforms. in 2022, mareano is also starting work in the north sea. mareano produces a range of thematic map products. geological maps (e.g. bellec et al. 2017) include seabed substrate maps such as acoustic backscatter, grain size, sedimentary environment and genesis (fig.  1), intended for use at regional scales of 5°55´e 6°0´e 6°5´e seabed sediments (grain size) sandy mud muddy sand sand gravelly sandy mud gravelly muddy sand gravelly sand sandy gravel gravel sand, gravel and cobbles sand, gravel, cobbles and boulders gravel and cobbles gravel cobbles and boulders cobbles and boulders mud and sand with gravel, cobbles and boulders thin or discontinuous sediment cover on bedrock exposed bedrock 5°55´e 62 °3 0´ e 62 °3 0´ e 6°0´e 2 km 6°5´e fig. 1 example of detailed (1:20 000) seabed sediments (grain size) map from the southern part of nordre sunnmøre, one of the pilot areas for marine base maps for the coastal zone. background image: norge i bilder wms from www.geonorge.no. https://doi.org/10.34194/geusb.v52.8325 http://www.geusbulletin.org http://www.geonorge.no dolan et al. 2022: geus bulletin 52. 8325. https://doi.org/10.34194/geusb.v52.8325 4 of 16 www.geusbul let in.org 1:100  000–1:250  000, depending on the volume and quality of survey data in each area. geomorphological maps show the distribution of marine landscapes and landforms, including cold-water coral mounds predicted from topography. a suite of maps depicting geochemical composition and characteristics are also produced. the methods used for the production of these map products are outlined by bøe et al. (2020). the geological maps are used further to develop benthic habitat maps, which fuse biological, geological and oceanographic information. mareano results are disseminated free of charge through www.mareano.no, www.ngu.no and many other portals, and data and results can also be obtained through www.geonorge.no. the multidisciplinary maps and data generated by mareano contribute to the scientific knowledge base for national ocean management plans. they are used widely by management institutions, petroleum and fisheries industries, academia and the public. 2.2 marine base maps for the coastal zone from 2020 to 2022, the norwegian mapping authority, ngu and imr are also collaborating on another coordinated pilot project for seabed mapping and data distribution: marine base maps for the coastal zone. the methodology is partly based on previous marine basemap projects by ngu (elvenes et al. 2019; bøe et al. 2020) and draws on experience from mareano as well as previous work of the partner institutions. the project produces a range of hydrography, geology, biology, nature type and geochemistry maps in three example areas to lay the groundwork for a national coastal mapping programme. a proposal for extending the pilot to a national programme, which will map approximately 100  000 km2 of seabed in coastal norway, was delivered to the norwegian government in 2021 with a view to starting in 2023. this mapping is estimated to take 15–20 years within the proposed framework at a cost of around 4 billion nok (c. 400 million eur). amongst the most notable conclusions from a recently conducted socio-economic analysis of this proposal, the financial investment is very profitable. furthermore, when the mapping is complete, the invested amount can be saved every year due to cost savings and smart decisions informed by comprehensive baseline knowledge of the seabed. in other words, the mapping will pay for itself over time and would mean a financial loss if not conducted. geological thematic maps from marine base maps for the coastal zone are like those produced in mareano spanning geology (grain size, sedimentary environment and genesis; fig. 1) and geomorphology (landforms). they are produced at a finer scale of 1:20 000 through expert interpretation of all available data providing the necessary higher level of detail for coastal management. additionally, several applied thematic maps including anchoring conditions, diggability and accumulation basins have been developed from the main geological map products. surface sediment samples and short cores from accumulation basins are analysed for organic and inorganic components to evaluate levels of contamination and temporal evolution (past 100–200 years). the geological maps are used further within the project to produce nature-type maps according to nature in norway (nin; see section 2.3). the geological information is used alongside hydrographic and oceanographic variables to provide environmental predictor variables, which are combined with classified observations of major and minor nature types (determined from biological and environmental characteristics) from video data. all geological maps are published online at www. ngu.no and are also available via other national and international portals, including www.mareano.no. some examples of geological map products are shown in figs. 1 and 2. other multidisciplinary results from the pilot project are available from the norwegian mapping authority and imr. whilst the pilot project is active, they are also available through the dedicated portal marinegrunnkart.avinet.no. 2.3 other initiatives in norway relevant to seabed geodiversity scandinavian countries have generally been recognised as forerunners in the promotion of geodiversity, with the work of johansson (2001), which highlights the (terrestrial) geodiversity of the region, being highly praised by brilha et al. (2018). in norway, the nature diversity act (ministry of climate and environment 2009) came into force in 2009. this act aims to promote conservation and sustainable use of the ‘full range of variation of habitats and landscape types’. it means that geological and landscape diversity have been officially recognised alongside biological diversity, but also that information on their spatial distribution must exist for successful implementation of the act. the need for this type of information began to be addressed through initiatives such as the norwegian programme for mapping of marine habitats (bekkby et al. 2011), which ran from 2007 to 2019. this national programme was designed to provide information on nature types selected under dn handbook 19 (direktoratet for naturforvaltning 2007). this programme included aspects of geological mapping, geological features and geodiversity, with ngu contributing maps of carbonate (shell) sand occurrences and ice marginal deposits. the need for information related to the nature diversity act also led to the development of nin, to provide a unified framework for delivering the required knowledge. the nin framework facilitates classification and description of nature across terrestrial, freshwater and https://doi.org/10.34194/geusb.v52.8325 http://www.geusbulletin.org http://www.mareano.no http://www.ngu.no http://www.geonorge.no http://www.ngu.no http://www.ngu.no http://www.mareano.no http://marinegrunnkart.avinet.no dolan et al. 2022: geus bulletin 52. 8325. https://doi.org/10.34194/geusb.v52.8325 5 of 16 www.geusbul let in.org marine systems and has become the national standard for publicly funded mapping initiatives. whilst the majority of the nin documentation is in norwegian (halvorsen et al. 2016), the approach was recently summarised for the international scientific community by halvorsen et al. (2020), who present nin as an implementation of the broader ‘ecosyst’ framework. nin includes several mechanisms for describing geological and geomorphological attributes, which are intended to feed into various proposed evs, including egvs. whilst nin has provided many opportunities and raised the bar for nature-type mapping, there remains scope for further development and practical testing of the system, especially in marine environments. erikstad (2013) discussed geodiversity as a comprehensive framework for management and conservation issues with special reference to norway, citing nin and similar approaches as part of the solution. through this discussion, the importance of maintaining geodiversity as a descriptor rather than merging it with management value is highlighted. this distinction is important across all aspects of nin and extends nicely to geological mapping within the ongoing projects we focus on here. 3. egvs – a new concept for seabed mapping? the egv concept (schrodt et al. 2019) highlights the need for consistency and proposes four main classes thin or discontinuous sediment cover on bedrock bioclastic sediment mass movement deposit and hemipelagic sediments mass movement deposit, locally covered by younger sediments debris flow deposit till, unspecified bedload (traction) deposit suspension deposit seabed sediments (genesis) landforms glaciotectonic hole submarine fan submarine slide canyon sediment wave field pockmark area glacial lineation ice-marginal moraine ridge, unspecified slide scarp channel current channel 18°0´e 18°0´e 17°0´e 17°0´e 16°0´e 16°0´e 70 °0 ´n 70 °0 ´n 10 km fig. 2 example of regional scale (1:100 000) seabed sediments (genesis) map produced for mareano at malangsgrunnen and surrounding area. interpreted landforms are overlain. https://doi.org/10.34194/geusb.v52.8325 http://www.geusbulletin.org dolan et al. 2022: geus bulletin 52. 8325. https://doi.org/10.34194/geusb.v52.8325 6 of 16 www.geusbul let in.org for organising geodiversity information: geology, geomorphology, soils and hydrology. many of the staple products of geological surveys and related institutions around the world are easily represented under these themes. however, these traditional products may be overlooked by users who perceive their interest to be focussed on other aspects of nature, for example, biodiversity. by elevating geological information to the same level as other evs through egvs, schrodt et al. (2019) help make geological information more visible in key global management arenas such as the united nations sustainable development goals. here, we examine the extent to which information relating to each egv class is already provided by ngu’s marine geological mapping, through initiatives like mareano and marine base maps for the coastal zone and evaluate scope for further development. in table 1, we have organised existing marine geological map products from ngu under the egv framework to assess the extent to which we are already mapping egvs. the egv classes most relevant to seabed geology are ‘geology and geomorphology’. under the egv class ‘geology’, ngu’s seabed map products fit naturally under ‘unconsolidated deposits’, with multiple map products providing a range of information per class. under current seabed-mapping initiatives, we do not specifically produce maps relating to ‘hardrock, fossil & mineral distribution’ or ‘geophysical processes’, although some existing map products from ngu are somewhat related to these topics. under the ‘geomorphology’ egv class, we deliver several products, which fit well as ‘landform distribution’ egvs. we have indicated the management or policy relevance for each of the egv classes based on those that ngu currently map for seabed-mapping initiatives. there is potential for developing additional, geodiversity-relevant map products related to several egvs, which will be discussed further in section 3.1. several maps planned for future production by ngu (e.g. sediment thickness maps, volume of sand or gravel deposits) will also provide additional relevant map products. the egv classes for ‘soils’ and ‘hydrology’, proposed by schrodt et al. (2019), are less relevant to marine geology and are not included in table 1. soil properties that would be important on land are generally absent and the relevant geological attributes, for example, chemistry and physical state, are captured under ‘geology – unconsolidated deposits’. hydrology too is largely irrelevant in the marine environment, except for submarine artesian wells, which are not mapped by ngu. according to schrodt et al. (2019), egvs should be (1) relevant to natural resource management and human well-being, conservation or ecology, (2) complementary to the other suites of evs and (3) feasible and cost effective to measure. generally, all our seabed maps relating to geology and geomorphology fulfil criteria (1) and (2). this is the very reason they are produced by ngu and are essential outputs funded through the major seabed-mapping initiatives. these geological map products form the basis for several applied or derived map products tailored specifically for various users and for nature-type mapping and management of the ocean areas. whilst the egv classes proposed by schrodt et al. (2019) are comprehensive, we also note scope for including an anthropogenic egv class. such information may be adequately highlighted separately on land and may have been deliberately omitted as an egv class for this reason or because it is not purely related to geodiversity. however, in the marine realm, which can tend to be ‘out of sight, out of mind’, there may be benefits to including this class under the egv umbrella. we have added some example features of geological relevance for further consideration, which are particularly relevant to egv criterion (1) in the seabed mapping. criterion (3) takes on a new meaning in the marine realm as compared to land. despite recent technical advances, seabed mapping is still an expensive exercise, particularly offshore, where large research vessels are needed. approximately 1.4 billion nok (c. 135 million eur) have been invested in mareano (offshore) at a cost of c. 5000 nok/km2 – nearly 500 eur/km2 (all products included, i.e. hydrography, geology, biology, habitats and chemistry), whilst estimates for the proposed coastal mapping programme are around 50 000 nok/ km2 – nearly 5000 eur/km2. it should be noted that the coastal mapping delivers a much larger suite of map products and is far more detailed than the mareano mapping (scale 1:20 000 versus 1:100 000 and coarser) and that multibeam surveys are considerably more costly in shallow waters. on land, whilst detailed geological mapping still requires considerable effort and fieldwork, a lot of first-pass geodiversity information can now be gained from satellite data and other relatively low-cost remote sensing and observations. by contrast, most of the seabed remains hidden from satellite imagery, requiring acoustic or other remote sensing methods to map the underwater topography and other acoustic indicators of seabed geological attributes (e.g. multibeam backscatter and sub-bottom profiler data). this generally requires access to suitable boats or other survey platforms which are expensive, especially in offshore and arctic (ice-influenced) waters. the mapping methods vary in their efficiency, cost-effectiveness and feasibility by water depth and practical considerations. furthermore, because acoustic remote sensing provides only a proxy to seabed geology, https://doi.org/10.34194/geusb.v52.8325 http://www.geusbulletin.org dolan et al. 2022: geus bulletin 52. 8325. https://doi.org/10.34194/geusb.v52.8325 7 of 16 www.geusbul let in.org table 1 comparison of essential geodiversity variables (egvs; schrodt et al. 2019) with existing marine geological map products from the geological survey of norway (ngu). original (from schrodt et al. 2019) marine (this paper) egv class1 egv definition examples seabed examples2 marine policy or management relevance hardrock, fossil and mineral distribution geological materials and their spatial distribution natural resources (e.g. coal, gas and ore) ngu bedrock maps cover land and some sea areas3 sustainable management spatial planning blue growth green transition nature conservation/ protection pollution management geohazard assessment climate change geology unconsolidated deposits surface distribution of parent materials resulting from geomorphological processes distribution or scarcity of materials (e.g. sand). dynamics of surface materials (e.g. sedimentation). standard products: seabed sediments (grain size) seabed sediments (genesis) sedimentary environment (present day erosion/deposition areas) accumulation of organic carbon sedimentation rates maps of chemical elements and compounds (organic and inorganic) bioclastic sediments (offshore) likely occurrences of coral reefs (offshore) shell-sand deposits (coastal areas) chemistry/pollution gas seeps applied/derived map products: sediment fractions sand and gravel resources anchoring conditions diggability accumulation basins submarine slides nin-specific environmental variables translated maps (e.g. emodnet classes) geophysical processes variability of the intensity of geophysical processes earthquakes volcanic eruptions earth radioactivity thermal energy land subsidence no specific maps4 geomorphology landform distribution landforms and their spatial distribution distribution of landforms resulting from erosion, transport and sedimentation dynamics of geohazards landforms (geomorphology) marine landscapes (physiographic regions) submarine landslides translated maps (e.g. emodnet landforms) sustainable management spatial planning nature conservation/ protection geohazard assessment human influence5 anthropogenic – – information on several human activities, e.g. dumping, dredging and trenching are currently included in the seabed sediment maps (grain size and genesis) but could be extracted as separate themes human impact nature conservation/ protection pollution management 1 egv classes soil and hydrology proposed by schrodt et al. (2019) are not included in this table. 2based on ngu maps produced for mareano/ marine base maps for the coastal zone. maps are available from https://www.ngu.no/en/topic/map-viewers, with options for download and map services also available. 3 further development of these products is not part of current seabed-mapping initiatives. 4 some ngu products are related to geophysical processes but fall more naturally under unconsolidated deposits or geomorphology. 5 potential ‘human influence’ egv class added, which is not part of schrodt et al.’s original egv classes. https://doi.org/10.34194/geusb.v52.8325 http://www.geusbulletin.org https://www.ngu.no/en/topic/map-viewers dolan et al. 2022: geus bulletin 52. 8325. https://doi.org/10.34194/geusb.v52.8325 8 of 16 www.geusbul let in.org it must be backed up by ground-truth data such as video and physical samples. the number of observations required is linked to the complexity of the area and the level of mapping detail required (spatial and thematic). this is generally a trade-off between available funds, information requirements and practical considerations. in the case of mareano and marine base maps for the coastal zone, a further consideration is that the ground-truthing campaigns are designed not only to meet the needs of geological mapping (verifying backscatter signatures, observing topographic features, etc.) but also to provide information for biological and habitat mapping as well as geochemistry. this multidisciplinary approach facilitates cost-effective use of resources and paves the way for multiple map products spanning many evs, as well as collaboration on products of common interest, for example, habitat maps and geochemistry, ultimately delivering a comprehensive suite of information for management and other users. the multidisciplinary mapping approach inevitably leads to compromises as compared to single-objective mapping, but overall, our experience is that the benefits (also economic) far outweigh the limitations (see also section 4). 3.1 highlighting geodiversity using egvs besides seeing how well our products fit into the egv classes and provide a basis for geodiversity related information, we are interested in how far the egv framework goes towards highlighting geodiversity per se. at face value, the framework seems to highlight relevant information from which one can gain insight into geodiversity through geological mapping but is not prescriptive as to how to provide geodiversity information directly in the form of (semi-) quantitative indices. we are mindful of comments by erikstad (2013) who, in discussing the need to measure diversity, points out that simple solutions such as counts of different units can be problematic due to oversimplification. this may be particularly concerning if these metrics are used outside the original intended context and used to associate value. we note that the egv concepts appear to build on earlier work by several of the authors who contributed to schrodt et al.’s (2019) paper where the term ‘geodiversity components’ (gdcs) is used to refer to quantified geofeatures ( bailey et al. 2017) across similar themes (geofeature categories) via calculations of coverage, richness or other dimensions. gdcs are inherently tied to the scale at which these various measures of geodiversity are studied. calculations applicable to midto broad-scale geodiversity in the context of land management are discussed by pellitero et al. (2015) who summarise many of the approaches reported in previous literature. one of the most common methods measures is richness per unit area, which has also been used in marine geological applications (e.g. kaskela & kotilainen 2017) for various components of their geodiversity assessment. through various projects, ngu has recognised the need to adapt traditional geological map products to end users and for onward use. several examples of applied maps derived from categorical maps of surficial sediments are reported by elvenes et al. (2019; e.g. anchoring conditions, diggability, etc.). whilst none of these specifically highlight or quantify geodiversity yet, they, nevertheless, demonstrate a need to translate geological information into more readily digestible formats for a variety of end users. figure 3 summarises some sediment grain size (categorical) grain-size fractions nin variables, e.g. categories of finegrained material applied products, e.g. diggability translation to other categories, e.g. emodnet geology substrate diversity additional product development, e.g. habitat modelling fig. 3 conceptual diagram showing how applied map products may be developed from a traditional marine geological map (categorical sediment grain-size map). these applied products may be better suited to onward use for various purposes. they illustrate one way in which we may make essential geodiversity variables (egvs) even more relevant for sustainable development. https://doi.org/10.34194/geusb.v52.8325 http://www.geusbulletin.org dolan et al. 2022: geus bulletin 52. 8325. https://doi.org/10.34194/geusb.v52.8325 9 of 16 www.geusbul let in.org ways in which a traditional sediment grain-size map may be converted into alternative forms to suit a variety of purposes. whilst we recognise potential risks of overgeneralising geological information by reducing it to estimates of component parts of geodiversity (or gdcs), we also recognise a perhaps greater risk of not making it accessible enough by letting it remain hidden in traditional map products, which may not match the needs of an expanding suite of end users. the development of additional map products is an ongoing process at ngu and has recently included several that specifically target geodiversity. here, we preview two of these maps, which highlight substrate diversity and morphological diversity. 3.1.1 substrate diversity the difference in inherent grain-size diversity between grain-size categories is well known to a geologist familiar with a given system of grain-size classes, such as those used by ngu (bøe et al. 2010; bellec et al. 2017; elvenes et al. 2019; geological survey of norway 2022). however, the differences in relative composition, and, hence, the likely diversity of substrates available within each of the different sediment polygons, may not be clear to all end users. similarly, the spatial variation of a particular grainsize fraction (e.g. mud content) may be obscured by traditional classifications and map symbology. conversion of the original categories to a suite of maps showing the spatial distribution of constituent fractions can be invaluable for many applications, aiding onward use of the geological maps. for example, van son et al. (2020) used the proportion of hard substrate estimated from ngu’s maps to model the spatial distribution of kelp biomass. this work used interpolated estimates of the sediment-fraction distribution (using kriging); however, we have found that for many applications, including mareano biotope modelling, a simple translation per polygon is sufficient. a lookup table listing the fractional content per class is provided for reference (see supplementary file s1) including all classes used to date in ngu maps; this updates the version used in van son et al.’s (2020) study. alternatively, we may translate categorical grainsize maps into other categorical schemes for onward use, for example, nin fine-material content classes (norwegian biodiversity information centre 2022), or to facilitate harmonisation with international data, for example, emodnet geology substrate classes (kaskela et al. 2019; vallius et al. 2020). conversion of traditional maps to new formats also provides opportunities for highlighting relationships between sediment properties and aspects of geodiversity, which may, otherwise, be rather hidden. for example, following conversion of our categorical sediment map to component fractions, we can quantify substrate diversity by calculating the entropy between layers. although often applied to probability layers, for example, for quantifying between class uncertainty in habitat (dolan et al. 2021) or soil mapping (hengl et al. 2017), the entropy method is generic and can be applied to our estimates of sediment fractions derived from the categorical sediment map. figure 4 illustrates how ngu’s categorical grain-size maps can be translated into constituent fractions and used to produce a map of substrate diversity. since shannon entropy has previously been applied as a measure of geodiversity between adjacent units within a neighbourhood radius (e.g. read et al. 2020), we emphasise that here we are determining the entropy of the values of co-located pixels between our five overlapping raster layers representing the fractions. whilst a neighbourhood version could theoretically be applied to a categorical sediment map, it would fail to yield meaningful information on substrate diversity since the classes themselves have intrinsic diversity. such an analysis would thereby only yield class diversity, highlighting transition zones between categories. 3.1.2 morphological diversity similar concepts apply to landforms and other geomorphic features, and it is important to recognise that landform diversity, whilst often linked to substrate diversity, may also be independent. for example, where landforms are covered by recent deposits of fine material, or where landforms (or bedforms) occur at finer scales than the changes in sediment properties, for example, sandwave fields. the mapping of landforms is often rather selective, based on project demands, mapping traditions, map scale and the possibility of reliable interpretation from available data and observations. the degree of generalisation of landforms is often not standardised from project to project, nationally and even less so internationally, making harmonisation challenging, for example, for emodnet (vallius et al. 2020). to help standardise geomorphological mapping, ngu has recently contributed to a two-part classification system for geomorphological features led by the british geological survey (dove et al. 2016, 2020), which separates morphological classification (part 1) from geomorphological interpretation (part 2). the applied study by nanson et al. (2022) further highlights the many benefits to such an approach, which we will not repeat in detail here. however, we note how the two-part approach lends itself naturally to opportunities to map morphological diversity separately from the diversity of geomorphic features, which are tied to specific geological processes. ngu is working to further develop this two-part approach with particular focus on glacial landforms. additional work aims to align the two-part classification with the recently revised landform list for nin (part of the nin descriptive system), which lists landforms by https://doi.org/10.34194/geusb.v52.8325 http://www.geusbulletin.org dolan et al. 2022: geus bulletin 52. 8325. https://doi.org/10.34194/geusb.v52.8325 10 of 16 www.geusbul let in.org geological process and will be published in the next version of nin (expected 2023). this is a far more comprehensive landform list than that currently implemented in nin or by ngu and provides a solid foundation for further work. morphological diversity is included in the gdcs used by bailey et al. (2017) who employed geomorphons (jasiewicz & stepinski 2013) to map landform units. the geomorphon approach uses computer vision techniques to analyse raster digital terrain models and is well suited to automated mapping of morphometric units, offering several advantages of earlier approaches (e.g. fisher et al. 2004). bailey et al. (2017) used the geomorphon classes to obtain coverage estimates of each morphometric type per unit area as a predictor of biodiversity. coverage estimates are a perfectly viable form of geodiversity metric for seabed mapping too, but we argue that a more intuitive impression of the spatial variation of morphometric diversity can be gained by using morphometric richness estimates at spatial scales relevant for various applications. using the bress toolbox (masetti et al. 2018), which presents an implementation of the geomorphon method specifically targeted to bathymetry (and optionally reflectivity) data, we have explored the potential of this method for mapping morphometric features in a variety of geological settings. the results are encouraging, across a range of bathymetric data resolutions, and using various options for the number of morphometric classes. sowers et al. (2020) recently used bress to show how geomorphons can be used to classify coarse (100 m) bathymetry data to help map broad scale geomorphology linked to the united states coastal and marine ecological classification standard (cmecs; federal geographic data committee 2012). in fig. 5, we show how geomorphon classification can be applied to data of different resolutions to gain a nested impression of morphometric features. this may aid expert interpretation of landforms, or as is our focus here, these classifications may be used to compute geomorphon richness (and optionally patchiness). this may allow us to highlight morphological diversity in a more complete way than via traditional (selective) landform mapping (fig. 5a). our example includes geomorphon classification applied to two resolutions of the same data. we use the 10-class option as per the original method of jasiewicz & stepinski (2013) and employed by bailey et al. (2017), though other options are available via bress. geomorphon settings (inner radius 50 m, outer radius 200 m, 50 km original categorical sediment grain-size map substrate-diversity map – entropy of 5 fractionsfive fractions mud sand gravel cobbles & boulders rock 50 km 73 °0 ´n 73 °0 ´n 74 °0 ´n 74 °0 ´n 34°0´e32°0´e 38°0´e36°0´e34°0´e fig. 4 example showing how geological survey of norway (ngu)’s sediment grain size map may be converted to five component fractions (mud, sand, gravel, cobbles, and boulders and rock), where darker colours indicate higher percentage content. note that in this figure, rock is only non-zero within the black box where it occurs very locally (few pixels only). a preliminary substrate diversity map is computed from these fractions using the entropy between the five fractions; darker colours indicate higher diversity. in the sediment grain-size map, blue shades indicate mud-rich sediments, yellow indicates sandy sediments and green indicates coarser mixed sediments. the full legend is available at www.ngu.no/mareano/grainsize.html and in supplementary file s1. https://doi.org/10.34194/geusb.v52.8325 http://www.geusbulletin.org http://www.ngu.no/mareano/grainsize.html dolan et al. 2022: geus bulletin 52. 8325. https://doi.org/10.34194/geusb.v52.8325 11 of 16 www.geusbul let in.org other settings default) were first used to capture local features visible in the 5 m bathymetry data (fig.  5b). second, we applied a broader scale geomorphon classification to 50 m bathymetry data (resampled from the 5 m data using bilinear resampling). here, we used the same length scales as used by sowers et al. (2020; inner radius 300 m, outer radius 1500 m), adapted for a 50 m bathymetry grid, with other settings as default (i.e. no adaptation of the flatness parameter, as adjusted by sowers et al. (2020)). examining fig. 5b (and inset), we see that fine-scale features, including iceberg ploughmarks and pockmarks, are effectively captured by the geomorphon analysis using local settings. larger features, including the prominent glacial meltwater b c d e * * * * * * a pockmark area area with ridges grounded iceberg depression hill (glaciotectonic) glaciotectonic hole ice marginal moraine drumlin slide front slide scarp ridge, unspecified drumlin esker glacial lineation glacial meltwater channel landforms 74 °0 ´n 73 °0 ´n pit valley footslope concave slope slope convex slope shoulder ridge peak flat geomorphons 101 geomorphon richness (km²) 33°0´e 31°0´e 34°0´e 32°0´e 35°0´e 33°0´e 36°0´e 34°0´e 37°0´e 35°0´e fig. 5 example area in the barents sea mapped by mareano, showing traditional landform mapping alongside morphometric feature classification using geomorphons at fine and broad scales. (a) standard geological survey of norway (ngu) landform map (offshore), which includes a combination of polygon and line features mapped by expert interpretation of bathymetric and supporting geological data. (b) fine-scale geomorphon classification of 5 m bathymetry data. (c) geomorphon richness showing the number of fine-scale geomorphon classes (from b) per km2 (5 m grid). (d) broad-scale geomorphon classification of 50 m bathymetry data. (e) geomorphon richness showing the number of broad-scale geomorphon classes (from d) per km2 (50 m grid). https://doi.org/10.34194/geusb.v52.8325 http://www.geusbulletin.org dolan et al. 2022: geus bulletin 52. 8325. https://doi.org/10.34194/geusb.v52.8325 12 of 16 www.geusbul let in.org channels, are delineated by the broader scale analysis shown in fig. 5d. the interpreted landform map (fig. 5a), by contrast, includes selected features spanning these two scales. it links mappable morphological units to the geological process that created them, a step that requires expert interpretation and supporting data sets, for example, sub-bottom profiler data. only those landforms in which mareano and ngu maps as standard are included, and only where sufficient information exists to determine their origin. note that individual pockmarks are not mapped due to their immense numbers (rise et al. 2014), but large areas of pockmarks are delineated in ngu maps as polygon features. also, iceberg ploughmarks are not delineated in ngu maps since they are so widespread, diverse in form and scale, and often overlapping (e.g. bjarnadóttir et al. 2016). whilst the landform map gives invaluable geological information, it does not provide such a good basis for assessing morphological diversity as the morphometric features mapped using geomorphons, which can also serve as a valuable complementary map product. the variety of morphometric features (geomorphon classes) within a given analysis neighbourhood can provide a measure of morphological diversity (richness). we use the explicit term ‘geomorphon richness’ here to avoid confusion. several studies, across various applications, have used focal statistics for similar diversity estimations; however, our testing confirms that for focal analyses using larger neighbourhoods, the results become dominated by artefacts associated with the analysis window itself. this effect was noted by wilson et al. (2007) in relation to terrain attributes and is just as applicable here. whilst the roving window approach using focal statistics may be successful in some applications (e.g. kaskela & kotilainen 2017), the approach presented here is less computationally intensive or prone to artefacts, aiding analysis across many scales. we convert the bress geomorphon output to integer, applying a majority filter to the results before computing diversity (this eliminates single pixels of a particular class, which are generally of little practical value). by overlaying a fishnet at the scale of interest, we can then extract the variety of geomorphon classes within each grid cell using zonal statistics. here, we show results using a 1 km grid (figs 5c, e), which provides an overview of geomorphon diversity (richness) at the mesoscale, sensu greene et al. (1999) within the study area. for megascale analyses, such as a national level, summaries at a larger (e.g. 10 km) grid scale may be more suitable. the raster output from the zonal statistics can be set to the desired resolution for onward use. by default, the cell size is the same as the input raster. in norway, due to military restrictions on bathymetry data within the 12 nautical mile territorial boundary, 50 m analysis offers the minimum practical size for unified analysis of geomorphons on a national scale. this will fail to capture fine-scale features and, hence, their diversity. however, as we see from fig. 5e, many important larger features are still captured, and this approach can give a very good indication of morphometric diversity when applied to larger data sets. further development of this geomorphon-based approach is ongoing at ngu, including examining ways in which it may support landform mapping, but initial results are promising. results to date also suggest that the morphometric features detected using similar distance settings are consistent across data resolutions, depending on the information content of the data (i.e. coarse data cannot detect small features). in addition to providing a basis for diversity assessments, geomorphon-based morphometric classification provides a useful complement to the interpreted landform map. it may be used for onward product development, for example, habitat mapping (wyles et al. 2022), or to aid the planning of ground-truthing cruises. where we find high morphometric diversity and other environmental diversity, we may expect greater biodiversity and may, therefore, require greater sampling effort (van son et al. 2015). 4. discussion 4.1 multidisciplinary mapping and egvs in this section, we reflect on how our multidisciplinary mapping programmes have helped map norway’s seabed geology and geomorphology, thus delivering egv-relevant information (table 1). ngu’s partnership with collaborating institutions (norwegian mapping authority and institute of marine research) has been invaluable in both mareano and the marine base maps for the coastal zone pilot. working under the same programme umbrella allows planning, execution and delivery of map products to be aligned, as well as meeting the needs of multiple end users. for instance, the geological mapping is very much dependent on multibeam echosounder data (bathymetry, backscatter and water column data). by partnering with the norwegian mapping authority, who has the overall responsibility for this data acquisition and bathymetric data processing, we gain better quality data, which can be used for multiple purposes. similarly, through the partnership with ngu and imr, the norwegian mapping authority (traditionally focussed on safety of navigation) gains insight into bathymetric data quality issues important for geological interpretation and use in habitat mapping, which may be irrelevant for hydrography. examples could be data artefacts that lead to misleading terrain attributes (lecours et al. 2017a, 2017b) or overenthusiastic data https://doi.org/10.34194/geusb.v52.8325 http://www.geusbulletin.org dolan et al. 2022: geus bulletin 52. 8325. https://doi.org/10.34194/geusb.v52.8325 13 of 16 www.geusbul let in.org cleaning, which obscures real morphological features in deeper waters. additionally, fuller use of backscatter and water-column data are made through partnerships. the backscatter data are an invaluable proxy to sediment type (given sufficient ground truthing) and fundamental to the development of good geological maps (which, in turn, feed into habitat and nature-type maps). backscatter-processing expertise and links to international initiatives for improving data quality, for example, geohab backscatter working group (lurton & lamarche 2015), are more accessible to the norwegian mapping authority, thanks to ngu’s involvement. water-column data are not only used to detect gas seeps, which themselves are an important nature type under nin, but also of interest for the oil and gas industry and important for the study of natural pollution and links to climate change (ruppel & kessler 2017). fieldwork is more efficient and resource-effective through multidisciplinary cruises rather than multiple surveys for different objectives. video surveys are jointly planned, and data are shared for geological and biological interpretation. likewise, some samples are shared to meet different objectives (e.g. shared grab sample for sediment ground truthing and infauna sampling; shared multicore deployments for organic and inorganic chemistry and microplastics) making efficient use of ship time. additional uses of the geological map products are more easily realised in a multidisciplinary framework, where partners are aware of and closely connected with each other’s work. geological and bathymetric maps are used directly in habitat and nature-type mapping and may be adapted to suit specific purposes, thanks to the collaboration. this has helped cement geology as an integral part of nature-type mapping and contribute to the development of nin, which, in turn, has benefits for nature conservation and management. there may be a perceived risk that multidisciplinary mapping could hinder geological mapping due to a diluted focus. in our experience, this is unfounded; instead, it has been advantageous to be able to combine many and different data sets and work across disciplines. there have been few compromises, and those made have been largely outweighed by opportunities for additional mapping or follow-up studies (including by universities and other institutions outside the core programme partnership). there are also numerous less tangible benefits related to exchange of ideas and a greater common understanding between partner institutions that have arisen from the multidisciplinary approach. it is difficult to imagine how a single-discipline approach can meet the demand for knowledge that will be needed for effective and sustainable management of our planet’s interconnected systems. it is equally difficult to imagine how evs for sustainable management in line with the united nation’s sustainable development goals can be effective without the inclusion of egvs alongside the other evs. in this paper, we have seen how geological and geomorphological maps deliver such information and potential for developing additional map products that highlight geodiversity. as norway and other nordic countries continue to produce a thorough suite of geological and geomorphological map products that fall neatly under the egv framework, we will see how egvs provide a good platform for geological mapping on the global environmental stage. with geological surveys and similar institutions seeking to gain greater relevance for their work, the egv ‘brand’ can be beneficial, just as we have witnessed with the development of nin. whether it can help secure funding for geological mapping is not a question we can answer, but we suggest it should be more of a help than a hinderance. 4.2 general comments on the egv framework there is still considerable scope for the development of the egv framework. it is broad, quite generic and is not tied to specific map scales or products. we have shown that it is relatively easy to adopt the framework in ‘well mapped’ areas of the seabed such as those we have presented from norway. here, geological map products (mostly 1:20  000–1:250  000) are based on a comprehensive suite of data, and as we have shown, several of these have the potential to be translated to additional products that highlight geodiversity. in areas with less complete data, either in terms of geographic coverage or information content, it may be more challenging to deliver useful egv-relevant information on sediment properties. the scales at which geological map products can be produced are tied to the availability of a suite of data, which is, in turn, linked to cost and access limitations. mapping generally combines remotely sensed data and ground-truth observations. where observations are sparse, and/ or remotely sensed data are coarse (pixel sizes of several hundred metres), the mapping scales may be in the order of 1:1 000 000 and coarser (e.g. 1:3 000 000 sediment grain-size map for the barents sea in lepland et al. (2014)). comparing this with mareano maps for the same region (e.g. via ngu’s online map service geo.ngu.no/kart/marin_mobil), we emphasise how generalised the sediment classes are, and we are uncertain to what extent a useful level of geodiversity information can be obtained from coarser-scale seabed mapping such as this. we note, however, that laverick et al. (2022) showed the potential for extending the suite of geological information from broad-scale maps (including lepland et al. 2014) using predictive modelling. despite their https://doi.org/10.34194/geusb.v52.8325 http://www.geusbulletin.org http://geo.ngu.no/kart/marin_mobil dolan et al. 2022: geus bulletin 52. 8325. https://doi.org/10.34194/geusb.v52.8325 14 of 16 www.geusbul let in.org coarse resolution and inherent uncertainty, broad-scale geological maps (e.g. lepland et al. 2014; diesing 2020; laverick et al. 2022) still provide invaluable information in areas where data are otherwise lacking and can help to prioritise follow-up studies. regarding seabed morphology, maps of classified morphometric features, interpreted landforms and landscapes (physiographic regions) are well matched to delivering egv-relevant information across the entire range of mapping scales, for which topographic and supporting data are available. here, we have shown how algorithms such as geomorphons may contribute to this effort and further lend themselves to quantitative assessments of diversity, alongside terrain attributes, for example, slope and relative relief, which provide complementary information. whilst more specific goals (including guidance on data requirements and mapping scales) for egv product development may be sought by some, the open scope allows the egv concept to be very inclusive. this is an important quality for global adoption. in our case, for mapping new and relatively inaccessible areas of the seabed, as well as those already subject to multiple user-group pressures, more complete links between egvs and gdcs for quantifying seabed geodiversity will help to make the geological information more tangible. here, we have shown some examples of how such product development might start, but there is a wealth of possibilities to be explored, including further links to biodiversity, which should be more easily realised through multidisciplinary mapping. 5. conclusions in this paper, we have provided an overview of the major seabed-mapping initiatives currently active in norway. we set out to assess the extent to which these initiatives deliver geodiversity information in relation to the egv framework. specifically: • we found that the existing map products deliver geological and geomorphological information that fits neatly within the egv framework’s geology and geomorphology classes. there are no obvious gaps in knowledge although there remains scope for development of additional products to meet more specific needs. several of these could be derived from existing maps without significant additional effort (e.g. translation to sediment fractions); others require additional data (e.g. sediment thickness maps). • we provided examples of how quantitative measures of geodiversity can be obtained from further development of existing products (substrate diversity) and through supplementary analysis of existing data (morphological diversity). • we have highlighted how multidisciplinary seabed-mapping initiatives help delivery of geodiversity information, which fits the egv framework by providing greater opportunities for effective mapping and knowledge development. the multidisciplinary approach helps cement geological information as part of an essential suite of information for sustainable management nationally. translating this to a global stage, our experience supports the argument for egvs standing alongside other evs to provide knowledge for long-term global sustainable management. acknowledgements we wish to thank all participants of the mareano programme and marine basemaps for the coastal zone (pilot) project. we thank the reviewers for constructive comments on the manuscript. multibeam bathymetry data shown in fig. 5 were acquired and supplied by the norwegian hydrographic service (kartverket) for mareano. the data are released under a creative commons attribution 4.0 international (cc by 4.0): https://creativecommons.org/licenses/by/4.0/. additional information funding statement this study was supported by the geological survey of norway and draws on results from the mareano programme and the marine base maps for the coastal zone pilot project. competing interests the authors declare no competing interests. author contributions mfjd: conceptualisation, methodology (quantitative diversity examples), writing – original draft. rb: project administration, writing – review and editing. lrb: project administration, writing – review and editing. additional files one additional file (supplementary file s1) is available at https://doi. org/10.22008/fk2/e0ipc9. references bailey, j.j., boyd, d.s., hjort, j., lavers, c.p. & field, r. 2017: modelling native and alien vascular plant species richness: at which scales is geodiversity most relevant? 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https://doi.org/10.3389/fmars.2020.00107 https://doi.org/10.1080/01490410701295962 https://doi.org/10.3389/fmars.2022.818635 https://doi.org/10.3389/fmars.2022.818635 delivering seabed geodiversity information through multidisciplinary mapping initiatives: experience 1. introduction 2. geodiversity as part of multidisciplinary seabed-mapping initiatives in norway 2.1 mareano 2.2 marine base maps for the coastal zone 2.3 other initiatives in norway relevant to seabed geodiversity 3. egvs a new concept for seabed mapping? 3.1 highlighting geodiversity using egvs 3.1.1 substrate diversity 3.1.2 morphological diversity 4. discussion 4.1 multidisciplinary mapping and egvs 4.2 general comments on the egv framework 5. conclusions acknowledgements additional information funding statement competing interests author contributions additional files references figures fig. 1 example of detailed (1:20 000) seabed sediments (grain size) map from the southern part of nordre sunnmøre, one of the pilot areas for marine base maps for the coastal zone. background image: norge i bilder wms from www.geonorge.no. fig. 2 example of regional scale (1:100 000) seabed sediments (genesis) map produced for mareano at malangsgrunnen and surrounding area. interpreted landforms are overlain. fig. 3 conceptual diagram showing how applied map products may be developed from a traditional marine geological map (categorical sediment grain-size map). these applied products may be better suited to onward use for various purposes. they illustrate one way in which we may make essential geodiversity variables (egvs) even more relevant for sustainable development. fig. 4 example showing how geological survey of norway (ngu)’s sediment grain size map may be converted to five component fractions (mud, sand, gravel, cobbles, and boulders and rock), where darker colours indicate higher percentage content. note that in this figure, rock is only non-zero within the black box where it occurs very locally (few pixels only). a preliminary substrate diversity map is computed from these fractions using the entropy between the five fractions; darker colours indicate higher diversity. in the sediment grain-size map, blue shades indicate mud-rich sediments, yellow indicates sandy sediments and green indicates coarser mixed sediments. the full legend is available at www.ngu.no/mareano/grainsize.html and in supplementary file s1. fig. 5 example area in the barents sea mapped by mareano, showing traditional landform mapping alongside morphometric feature classification using geomorphons at fine and broad scales. (a) standard geological survey of norway (ngu) landform map (offshore), which includes a combination of polygon and line features mapped by expert interpretation of bathymetric and supporting geological data. (b) fine-scale geomorphon classification of 5 m bathymetry data. (c) geomorphon richness showing the number of fine-scale geomorphon classes (from b) per km2 (5 m grid). (d) broad-scale geomorphon classification of 50 m bathymetry data. (e) geomorphon richness showing the number of broad-scale geomorphon classes (from d) per km2 (50 m grid). table table 1 comparison of essential geodiversity variables (egvs; schrodt et al. 2019) with existing marine geological map products from the geological survey of norway (ngu). research article | short dolan & bjarnadóttir 2023: geus bulletin 52. 8337. https://doi.org/10.34194/geusb.v52.8337 1 of 7 highlighting broad-scale morphometric diversity of the seabed using geomorphons margaret f.j. dolan* , lilja r. bjarnadóttir geological survey of norway (ngu), trondheim, norway abstract morphometric diversity is an important component of overall seabed geodiversity. automated methods for classification of morphometric features (ridges, peaks, valleys etc.) provide a convenient way of classifying large volumes of data in a consistent and repeatable way and a basis for assessing morphometric diversity. here, we apply ‘geomorphons’, a pattern recognition approach to morphometric feature classification, to 100 m resolution multibeam bathymetry data in the barents and norwegian seas, norway. the study area spans depths from a few metres to nearly 6000 m across several geological settings. ten unique morphometric features are delineated by the geomorphon analysis. from these results, we compute the variety of features per 10 km2. this simple ‘geomorphon richness’ measure highlights broad-scale morphometric diversity across the study area. we compare the richness results with terrain attributes and across physiographic regions. our results provide new regional insights, which together with more detailed information will help guide follow-up surveys as well as identifying diversity hotspots, which may require special management. *correspondence: margaret.dolan@ngu.no received: 28 nov 2022 revised: 12 june 2023 accepted: 24 july 2023 published: 06 sept 2023 keywords: morphometric features, geodiversity, geomorphons, seabed mapping, marine geology abbreviations: cmecs: united states coastal and marine ecological classification standard mareano: marine areal database for norwegian sea areas ngu: geological survey of norway (norges geologiske undersøkelse) vrm: vector ruggedness measure geus bulletin (eissn: 2597–2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: kaskela anu (gtk, finland) and verner brandbyge ernstsen (geus, denmark) reviewed by: derek sowers (ocean exploration trust, usa) and giuseppe masetti (danish geodata agency, denmark) funding: see page 6 competing interests: see page 6 additional files: none introduction geodiversity (gray 2004) has many facets, or geodiversity components (bailey et al. 2017), including the diversity of morphometric features (ridges, peaks, valleys etc.). these can be mapped from topographic data using single or combined terrain attributes derived from digital elevation or bathymetry data or by expert interpretation of shaded relief (macmillan & shary 2009; lecours et al. 2016). alternatively, they may be delineated using specifically designed feature detection algorithms (e.g. dikau 1989; wood 1996; jasiewicz & stepinski 2013). morphometric features identified by these methods are often called ‘landforms’. here, we avoid this term, which is widely used in a broader sense and which has proved difficult to define and adopt consistently even within the fields of geomorphology and geomorphometry (evans 2012). they are also frequently described as geomorphological classifications, but such features need to be put in context with their geological setting and/or the process(es) by which they have been formed for the purposes of geomorphological mapping. this important distinction is emphasised by dove et al. (2016, 2020) and nanson et al. (2022, 2023) in their two-part approach to geomorphological mapping, whereby morphometric features are mapped first, followed by geomorphological interpretation where viable. irrespective of the geomorphological origin of seabed morphometric features, their detection and classification using automated, algorithm-based methods offer a convenient starting point for estimating their diversity. numerous methods have been proposed for morphometric feature detection (macmillan & shary 2009). the pattern recognition approach of geomorphons (jasiewicz & stepinski 2013) is adopted here. this offers a more complete and computationally efficient classification than earlier approaches https://doi.org/10.34194/geusb.v52.8337 https://orcid.org/0000-0003-4405-9277 https://orcid.org/0000-0002-3152-8348 mailto:margaret.dolan@ngu.no dolan & bjarnadóttir 2023: geus bulletin 52. 8337. https://doi.org/10.34194/geusb.v52.8337 2 of 7 www.geusbulletin.org (e.g. dikau 1989; wood 1996; fisher et al. 2004; schmidt & andrew 2005), which need to be combined to capture a complete set of morphometric features and can struggle when applied to large data sets. the algorithm detects 498 geomorphons, which are reduced to the ten most frequent and commonly recognisable features (flat, peak, ridge, shoulder, spur, slope, hollow, footslope, valley and pit) through reclassification (jasiewicz & stepinski 2013). the number of classes may nevertheless be changed to match user needs via user-defined lookup tables (e.g. masetti et al. 2018; masetti 2022). the geomorphon approach has existed since 2013, but it has only recently seen uptake in the marine environment (e.g. dekavalla & argialas 2017; di stefano & mayer 2018; masetti et al. 2018; novaczek et al. 2019; sowers et al. 2020). geomorphons also appear to have only recently been explored as a basis for terrestrial geodiversity assessments (bailey et al. 2017; pál & albert 2021; vörös et al. 2021) and to our knowledge have not been applied in assessments of seabed geodiversity, besides the recent study by dolan et al. (2022). as well as contributing to a broader suite of geodiversity-related information, characterising morphometric diversity of the seabed has intrinsic value and can be particularly useful when investigating new areas. in norway, the offshore seabed mapping programme mareano (bøe et al. 2020) has focussed on the continental shelf and slope but has recently begun mapping the deep norwegian sea, adding to earlier bathymetric mapping, with a view to providing information for sustainable management. this deep-sea area is far from flat and featureless. much of the terrain is far more extreme than the previously mapped shelf and slope areas – notably the hills and mountains of the mid-atlantic ridge and depths of nearly 6000 m at molloydjupet (molloy deep) west of svalbard. before attempting any follow-up surveys to further characterise the seabed geology and habitats, it is important to appreciate just how morphometrically diverse the deep-sea terrain is. alongside more detailed studies of the terrain, it is useful to conduct broad-scale analyses, which facilitate comparison with more familiar areas on the shelf and slope, and even on land. classification of morphometric features and assessment of their variety gives an insight into the morphometric diversity of an area, which complements information from traditional terrain attributes (slope, ruggedness, relative relief etc.). a research challenge in classifying these features is to develop methods that are repeatable and can be applied to very large data sets without overreliance on (often subjective) expert interpretation. sowers et al. (2020) used geomorphons as a foundation for classifying geomorphological units. using the united states cmecs geoforms (federal geographic data committee 2012) at 100 m resolution bathymetry, they showed that useful results can be obtained even from relatively coarse data. here, we use geomorphons to extract morphometric features on the norwegian seabed, using bathymetry data at this same resolution. we explore whether this approach provides a useful basis for quantifying broad-scale morphometric diversity, through determination of geomorphon richness per unit area (10 km2). this is then compared with examples of terrain attributes associated with morphometric diversity. furthermore, we examine how geomorphon richness varies by physiographic region. methods bathymetry data multibeam bathymetry data were compiled and supplied by the norwegian mapping authority hydrographic service (kartverket) for the mareano programme. this compilation (august 2021) combines bathymetry data from the mareano programme and related surveys. for this broad-scale study of the norwegian and barents seas, the data were resampled to 100 m resolution using bilinear interpolation. the elevation void fill function in arcgis v.10.8.1 was used to fill gaps in bathymetry data coverage of up to 2 km to minimise morphometric feature detection artefacts. such gaps appear only sporadically in the data, usually confined to older surveys over steep terrain. geomorphon analysis geomorphon analysis was conducted using the bress toolbox (masetti et al. 2018; masetti 2022), a free standalone tool developed specifically for applying the method to bathymetry data. masetti et al. (2018) and jasiewicz & stepinski (2013) detail the geomorphon algorithm while masetti’s (2022) user manual describes practical use of the bress toolbox. here, we provide an overview of the main user-defined settings. bress classifies into the original ten, or optionally six (default), five or four morphometric features (see lookup tables in masetti 2022). we use ten classes; this maintains the discrimination between convex and concave slopes (termed “spur” and “hollow” by jasiewicz & stepinski 2013) and maximises the number of different features for our morphometric diversity assessments. the inner and outer radii of the search annulus determine how near and far the algorithm will ‘look’ in each of eight directions. these settings affect not only which morphometric features will be detected but also the impact of noise or artefacts in the bathymetry https://doi.org/10.34194/geusb.v52.8337 http://www.geusbulletin.org dolan & bjarnadóttir 2023: geus bulletin 52. 8337. https://doi.org/10.34194/geusb.v52.8337 3 of 7 www.geusbulletin.org data on the results. sowers et al. (2020) found that while the default bress values of inner/outer radii of 5/10 grid nodes work reasonably well using 100 m resolution data, radii of 3/15 grid nodes provided results that were better matched with those that would be obtained by expert delineation of deep-sea terrain (>200 m depth). in this study, we use this same outer radius but increase the inner radius to seven grid nodes to help reduce the effects of noise in much of our older multibeam data from the deep-sea area. we notice that other data resolutions can be accommodated by adapting the radii used, while maintaining the distances considered. the other main user-defined setting is the flatness threshold (see jasiewicz & stepinski 2013). sowers et al. (2020) showed that adapting this setting, and thereby what the algorithm considers flat, is useful for delineation of specific morphometric features. for this study, which focuses on using geomorphons as a basis for assessing morphometric diversity, we maintain a constant (default) flatness value for consistency across all areas. we recognise that this may lead to the underor over-estimation of morphometric features, depending on the terrain, but it serves the purpose for a preliminary analysis. for more detailed analysis, with a view to geomorphological mapping, adjustments to the flatness parameter would be prudent. geomorphon richness the diversity of morphometric features was assessed by calculating the variety of geomorphon classes within a unit area. this calculation of ‘geomorphon richness’ employs a fishnet polygon grid as the basis for zonal statistics calculations, reporting the variety of classes within each polygon (dolan et al. 2022). here, using arcgis 10.8.1 with spatial analyst, we used a c. 10 km2 fishnet (3200 × 3200 m), which suits broadscale assessment, and calculated the geomorphon richness of an integer version of the bress classification output, which had first been subject to a majority filter (8 neighbours). terrain attributes and marine landscapes to explore the complementary information given by geomorphon richness, we generated two examples of terrain attributes: relative relief and vrm (sappington et al. 2007). both are commonly associated with morphometric diversity. relative relief was generated for a standard 1 × 1 km area using focal statistics (range) in arcgis 10.8.1 spatial analyst, while vrm was calculated via the spatialeco package in r (evans 2020) for a 3 × 3 cell analysis neighbourhood. we also compare geomorphon richness to the ngu’s marine landscape map, providing a semi-automatic classification of the major physiographic regions (elvenes 2014). results figure 1 shows the geomorphon classification for the entire study area. inset maps highlight the features captured for several example areas with different morphologies in various geological settings. near molloydjupet (fig. 1a), slopes dominate the extreme terrain, but peaks, ridges, valleys and shoulders are also effectively captured, as well as a few flat areas, including at the deepest part. on the mid-atlantic ridge (fig. 1b), we see a complex morphology featuring all geomorphon classes. of note is the delineation of volcanic cones as peaks. figure 1c highlights an area of the deep norwegian sea area mapped by older multibeam surveys with noisy data. here, we see that besides real (larger) features, the algorithm has classified noisy data (corrugations visible in shaded relief) as morphometric features, despite our best efforts to find an appropriate setting for the inner radius. at the barents sea shelf edge (fig. 1d), we see how larger morphometric features are effectively detected, including the sopphola-steinbitryggen hill-hole pair, the shelf edge and variations in slope morphology. however, several less prominent features are classed as flat. this is a good example of an area where adjustment of the flatness parameter may allow detection of more features. alternatively, analysis of higher resolution data may complement these results. at malangsdjupet (fig. 1e), the neighbouring strandflat (crystalline bedrock), banks and continental slope, we see effective delineation of the relatively flat bank areas (malangsgrunnen and sveinsgrunnen) from the morphometrically diverse coastal and continental slope areas. figure 2 shows the geomorphon richness results. although quite a coarse resolution, this figure complements the information in fig. 1 by highlighting areas of high and low morphometric diversity. the inset figures (a–e) show the richness for each example area, and the results of the relative relief calculations are included for reference. here, we see that morphometric features are effectively captured both in areas of high and low relative relief. as we might anticipate from fig. 1, area b on the mid-atlantic ridge is the most morphometrically diverse. we also see that diversity is overestimated in area c due to the noisy data leading to a high geomorphon richness. elsewhere, the summaries reflect the features shown in fig. 1. most importantly, the richness map highlights regional differences in the morphometric diversity: a property that is not immediately apparent from the geomorphon classification or terrain attribute maps. we gain a new appreciation of https://doi.org/10.34194/geusb.v52.8337 http://www.geusbulletin.org dolan & bjarnadóttir 2023: geus bulletin 52. 8337. https://doi.org/10.34194/geusb.v52.8337 4 of 7 www.geusbulletin.org where diversity hotspots are, and how much more diverse they are than previously mapped areas. furthermore, where we see apparently high diversity in unexpected places (e.g. deep sea plain), we are alert to the reasons for this, such as poor data quality. as well as visually assessing the geomorphons and their richness, we examined how the morphometric diversity varies by landscape type. the results are summarised in fig. 3 where we show the mean geomorphon richness by landscape type, with mean relative relief and ruggedness for reference. fjords, the strandflat and marine hills and mountains exhibit the highest values of geomorphon richness. the two former types have moderate relative relief and ruggedness at the analysis scales used, while hills and mountains alongside canyons have the highest relative relief and slightly higher ruggedness values. ruggedness values are generally low and due to the data resolution will not capture local variations, which may be linked to the, often fractal, morphology. relative relief is a useful indicator of how extreme the terrain is but does not appear to have a direct link to geomorphon richness at the scales analysed. discussion and conclusions we have applied geomorphons to classify 100 m resolution multibeam bathymetry data over large parts of the barents and norwegian seas, norway. we used this result to calculate geomorphon richness – a measure of morphometric diversity based on the variety of geomorphon classes per 10 km2. this has allowed us to gain initial insight into morphometric diversity hotspots (e.g. mid-atlantic ridge; figs 1a, 2a), which may be linked to high biodiversity, as well as identifying areas where the diversity is overestimated due to poor data quality (e.g. figs 1c, 2c). our results will help prioritise follow-up surveys to document surficial geology and benthic habitats using sampling and video as well as highlight areas where high resolution acoustic surveys (which would allow detection of finer-scale morphometric features) fig. 1 geomorphon classification of the study area in the norwegian and barents seas. the geomorphon classes are shown as a semi-transparent layer over greyscale hillshade. inset maps (a–e) show details of the geomorphon classes at several example locations. multibeam bathymetry: kartverket. background bathymetry (blue shaded relief): gebco bathymetric compilation group (2019). abbreviated place names: mor: molloyryggen. mod: molloydjupet. sr: steinbitryggen: sh: sopphola. mg: malangsgrunnen. md: malangsdjupet. sg: sveinsgrunnen. 200 km geomorphons flat peak ridge shoulder convex slope slope concave slope footslope valley pit c d e b a 5km 5km 5km 5km 5km mg sg10 km md sr 10 km sh 10 km 10 km 10 km mod mor a b c d e 20°0'w 10°0'w 0°0' 10°0'e 20°0'e 30°0'e 40°0'e 80°0'n 75 °0 'n 70 °0 'n 65 °0 'n 0°0' 10°0'e 20°0'e 75 °0 'n 70 °0 'n 65 °0 'n 80°0'n https://doi.org/10.34194/geusb.v52.8337 http://www.geusbulletin.org dolan & bjarnadóttir 2023: geus bulletin 52. 8337. https://doi.org/10.34194/geusb.v52.8337 5 of 7 www.geusbulletin.org would be particularly useful. this type of information can help sustainable management of the seabed and aid the design of marine protected areas. it is important that geodiversity is included in such management and conservation efforts (schrodt et al. 2019; tukiainen & bailey 2022), especially as this information may be available well in advance of biological data. quantitative methods for highlighting geodiversity are currently sparse. to promote geological information on national and international stages like the un sustainable development goals, a concerted effort to remedy this is needed, particularly in the marine environment. we hope that preliminary results such as those presented here will help spur on such development and provide an initial basis for comparison with biological data and biodiversity estimates. here, we compared geomorphon richness with examples of terrain attributes intuitively associated with morphometric diversity (and hence often with biodiversity). this showed that high morphometric diversity can occur in regions of high and low relative relief and ruggedness, but that low morphometric diversity seems to be associated with low values of these terrain attributes. furthermore, we examined how geomorphon richness varies by marine landscape, confirming that fjords, the strandflat and marine hills and mountains exhibit the highest diversity. our analysis was limited to a single resolution, relatively coarse data set. this was a practical solution for analysis of a large area, encouraged by the useful results obtained by sowers et al. (2020) with data of the same resolution. following limited testing, we adopted fixed values for bress user-defined settings across the entire study area to gain a first impression of how successful the method might be. further fine-tuning or splitting of areas may help make the classification truer to the real morphometric features present, for example by further reducing the effects of noise and better highlighting less distinct features. since this is challenging for a large area, a more promising approach may be to fig. 2 geodiversity overview and detailed inset maps (areas a–e) showing geomorphon richness (left) and relative relief (right). the colour ramps are common for all areas. note that the maximum value for the colour ramp of relative relief has been limited to 500 m to aid visualisation. values of up to 1248 m (up to 948 m in area a and up to 1087 m in area b) occur within the data set, but extreme values are rare. a maximum–minimum stretch has been applied as opposed to standard deviation or another stretch that will distort perception of the range of values. multibeam bathymetry: kartverket. background bathymetry (blue shaded relief): gebco bathymetric compilation group (2019). 10 km 10 km 10 km 10 km relative relief (m) 500 (+) 10 km 0 10 km e 10 km d c 10 km b 10 km a 10 km geomorphon richness index 10 1 a b c d e 10°0'w 0°0' 10°0'e 20°0'e 30°0'e 40°0'e 80°0'n 75 °0 'n 70 °0 'n 65 °0 'n 0°0' 10°0'e 20°0'e 75 °0 'n 70 °0 'n 65 °0 'n 200 km https://doi.org/10.34194/geusb.v52.8337 http://www.geusbulletin.org dolan & bjarnadóttir 2023: geus bulletin 52. 8337. https://doi.org/10.34194/geusb.v52.8337 6 of 7 www.geusbulletin.org combine broad-scale analysis, such as that conducted here, with local analyses based on more detailed data. this local analysis could be completed for most of the norwegian continental shelf where multibeam bathymetry data are available at 5 m resolution or better. bress user-defined parameters may need to be modified (via additional testing) to best identify morphometric features of interest at this scale. data for deeper areas are generally available at 25 m or coarser, so there is considerable potential for a finer resolution analysis at the best available resolution, if the study area is split into manageable sections (informed by e.g. depth or slope change thresholds) or tiled for computation purposes. it is worth noting that this finer resolution analysis (as per dolan et al. 2022) will likely result in a higher diversity of geomorphons for many areas and by extension will result in changes to the geomorphon richness results. analysis of the best available resolution data is, however, required for geomorphological interpretation. we hope that incorporating algorithm-based detection of morphometric features will aid our geological interpretation moving forward and support a two-part approach to geomorphological mapping (dove et al. 2016, 2020; nanson et al. 2023). more accurate estimates of morphometric diversity would be an invaluable parallel output from these analyses, alongside metrics of other components of geodiversity, supported by additional data and interpretation. acknowledgements the authors would like to thank the norwegian mapping authority hydrographic service (kartverket) for  access to the multibeam bathymetry data used in these analyses. the bathymetry data are available under a creative commons attribution 4.0 international (cc by 4.0) license at dybdedata.no. funding statement this study was supported by the geological survey of norway and draws on results from the mareano programme. author contributions mfjd: conceptualization, methodology, writing – original draft; lrb: project administration, writing – review and editing competing interests the authors declare no competing interests. references bailey, j.j., boyd, d.s., hjort, j., lavers, c.p. & field, r. 2017: modelling native and alien vascular plant species richness: at which scales is geodiversity most relevant? 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https://doi.org/10.1016/s0166-2481(08)00009-3 https://doi.org/10.1016/s0166-2481(08)00009-3 https://www.hydroffice.org/bress/main https://doi.org/10.3390/geosciences8010014 https://doi.org/10.1016/j.margeo.2022.106731 https://doi.org/10.5281/zenodo.7804019 https://doi.org/10.5281/zenodo.7804019 https://doi.org/10.1371/journal.pone.0216792 https://doi.org/10.1371/journal.pone.0216792 https://doi.org/10.5194/egusphere-egu21-1363 https://doi.org/10.2193/2005-723 https://doi.org/10.1111/j.1475-4762.2005.00638.x https://doi.org/10.1073/pnas.1911799116 https://doi.org/10.1073/pnas.1911799116 https://doi.org/10.3389/fmars.2020.00009 https://doi.org/10.3389/fmars.2020.00009 https://doi.org/10.1111/cobi.14024 https://doi.org/10.3390/geosciences11020058 https://hdl.handle.net/2381/34503 https://hdl.handle.net/2381/34503 highlighting broad-scale morphometric diversity of the seabed using geomorphons introduction methods bathymetry data geomorphon analysis geomorphon richness terrain attributes and marine landscapes results discussion and conclusions acknowledgements funding statement author contributions competing interests references figures fig. 1 geomorphon classification of the study area in the norwegian and barents seas. the geomorphon classes are shown as a semi-transparent layer over greyscale hillshade. inset maps (a–e) show details of the geomorphon classes at several example locations. multibeam bathymetry: kartverket. background bathymetry (blue shaded relief): gebco bathymetric compilation group (2019). abbreviated place names: mor: molloyryggen. mod: molloydjupet. sr: steinbitryggen: sh: sopphola. mg: malangsgrunnen. md: malangsdjupet. sg: sveinsgrunnen. fig. 2 geodiversity overview and detailed inset maps (areas a–e) showing geomorphon richness (left) and relative relief (right). the colour ramps are common for all areas. note that the maximum value for the colour ramp of relative relief has been limited to 500 m to aid visualisation. values of up to 1248 m (up to 948 m in area a and up to 1087 m in area b) occur within the data set, but extreme values are rare. a maximum–minimum stretch has been applied as opposed to standard deviation or another stretch that will distort perception of the range of values. multibeam bathymetry: kartverket. background bathymetry (blue shaded relief): gebco bathymetric compilation group (2019). fig. 3 mean geomorphon richness per landscape type (error bars indicate standard deviation) and comparison with terrain attributes, relative relief (rr) and the vector ruggedness measure (vrm). here, we show landscape types extracted for the study area. the full map, which is based on various sources of bathymetry data, is available at https://www.mareano.no and https://www.ngu.no. for further details on the landscape types, see https://www.ngu.no/mareano/landscape.html and elvenes (2014). areas a–e are indicated for reference to figs 1 and 2. geological survey of denmark and greenland bulletin 15, 2008, 73-76 in the last three field seasons the geological survey of denmark and greenland (geus) has undertaken mapping in the south-eastern part of the nuuk region in southern west greenland, and here we present new zircon ages that help constrain the northern boundary of the tasiusarsuaq terrane. the archaean geology of the nuuk region is commonly interpreted as a tectonic collage assembled through lateral accretion and collision of oceanic and continental slivers and blocks (e.g. friend & nutman 2005). popular jargon de scribes these as terranes, bounded by faults or mylonite zones and characterised by rocks of contrasting origin on either side of their tectonic boundaries (coney et al. 1980). the isukasia and færingehavn terranes (figs 1, 2) are the oldest terranes at ≥3.75 ga, and extend from the outer part of godthåbsfjord in the south-west to the margin of the inland ice in the north-east, but they might not have a common geological history (friend & nutman 2005). the tre brødre terrane is mainly represented by the ikkatoq gneiss and occurs in close spatial relationship with the færingehavn terrane, and also as a pronounced thrust unit along the qarliit nunaat thrust between the færingehavn and tasiusarsuaq terranes (fig. 1; nutman et al. 1989). the terrane boundaries in the inner fjord region near the inland ice margin are less well constrained; the tre brødre terrane extends into the region from the south-west, the kapisilik terrane is defined from the northern and eastern part and borders the tasiusarsuaq terrane to the south and possibly to the east. the terrane accretion is believed to have taken place in two events. the first terrane accretion is defined from the northern part of the region, and possibly involves the isukasia, kapisilik and akia terranes. the thermal event stitching these terranes is dated to c. 2.99–2.95 ga (fig. 2; hanmer et al. 2002; friend & nutman 2005). the second accretion phase of the major continental blocks is believed to have occurred at around 2.725–2.71 ga. this second event is well described, and in cludes anatexis and emplacement of continental crust-derived granites, which are associated with contemporaneous metamorphism (friend et al. 1996). figure 2 outlines regional plutonic, metamorphic and su pracrustal events. individual terranes were formed during relatively short time periods with active geological processes of creation and recycling of continental crust, and most of the terranes follow a similar pattern of development. the first plutonic events consisted of primitive magmas and produced to na lite–trondhjemite–granodiorite (ttg) and dioritic gn eisses. younger, more evolved granitic magmas were often intruded simultaneously with high-grade metamorphism. this development may reflect a stabilisation of the individual terranes. 73 new zircon ages from the tasiusarsuaq terrane, southern west greenland tomas næraa and anders scherstén © geus, 2008. geological survey of denmark and greenland bulletin 15, 73–76. available at: www.geus.dk/publications/bull greenland tre brødre færingehavn terrane tasiusarsuaq terrane terrane ‘nunatak 1390’ nuuk serm ilik 50°52° 64° 25 km qarliit nunaat thrust bu kse fjo rde n go dt hå bs fjo rd ? ameralik akia terrane kapisilik terrane nunataarsuq meso–neoarchaean granites meso–neoarchaean orthogneiss and granitic rocks. dots: granulite facies supracrustal belts (undifferentiated) anorthosite-gabbro complexes qôrqut granite complex terrane boundary palaeoproterozoic fault structural trend line eoarchaean gneiss 499227 499228 499161 499221 sample location 514832 fig. 1. geological map of the southern nuuk region (modified from escher & pulvertaft 1995), with locations and numbers of samples discussed in this paper. the tasiusarsuaq terrane the tasiusarsuaq terrane is dominated by 2.92–2.84 ga tonalite and granodiorite gneisses (friend & nutman 2001; crowley 2002). the main regional metamorphism is of amphibolite facies grade, however, granuite facies or retrogressed granulite facies rocks are present in large areas (fig. 1). peak granulite facies conditions have been dated at 2.81–2.79 ga (pidgeon & kalsbeek 1978; crowley 2002). greenschist facies rocks have been observed on ‘nunatak 1390’, which we in the present article suggest is part of the tasiusarsuaq terrane. northern boundary of the tasiusarsuaq terrane the north-western boundary of the tasiusarsuaq terrane in the buksefjorden area (fig. 1) has been described in some detail. narrow mylonite zones define a boundary between granulite facies gneisses of the tasiusarsuaq terrane and prograde amphibolite facies gneisses of the tre brødre terrane. prograde amphiolite facies metamorphism dated at around 2.74–2.70 ga is presumably related to the terrane accretion and has not been recorded within the tasiusarsuaq terrane itself (crowley 2002), which was thrust upon the tre brødre terrane during orogenesis (nutman et al. 1989). the eastern extension of the northern tasiusarsuaq terrane boundary remains speculative. here we present zircon u-pb age data from six selected rock samples collected in the vicinity of the proposed eastern extension of the northern tasiusarsuaq terrane boundary. tasiusarsuaq tonalite a migmatised tonalite representative of the basement gneisses in the northern part of the terrane was collected for zircon upb age determination (fig. 1; sample 499221 in the survey numbering system). the rock contains lenses of amphibolite and has abundant migmatite veins. palaeosome was separated from neosome by sawing slabs of each, and both sub-samples were dated. the internal zircon textures are very similar in both palaeosome and neosome. zircon grains have complex internal textures often with dark shells separating the core from the rim (fig. 3a, b). the cores display igneous oscillatory zonation to homogeneous textures. the age data for the palaeosome are concordant within 10% for 63 spots (n = 65), and regress to an essentially zero age lower intercept. we therefore use the 207pb/206pb ratios, which yield an age of 2.868 ± 0.004 ga (fig. 4a; n = 62/65, ± 2σ, mswd = 1.5). the outliers are slightly younger, presumably due to ancient pb loss. the neosome data are concordant within 10% for 55 spots (n = 57; fig. 4b); close inspection of the 207pb/206pb age data indicates that two ages at c. 2.87 and c. 2.80 ga can be differentiated (fig. 4b). however, this is speculative as the data suffer from insufficient precision in conjunction with an apparent ‘age smear’ as noted in the palaeosome, presumably due to ancient pb-loss. nevertheless, the two suggested ages at c. 2.87 and c. 2.80 ga are in excellent agreement with the palaeosome date and known ages for granulite facies metamorphism in the region (pidgeon & kalsbeek 1978; crowley 2002). discordant granite sheets 2.72 ga old cut the gneisses in this part of the tasiusarsuaq terrane (friend et al. 1996), but no such age component was found in this rock. ‘nunatak 1390’ and nunataarsuk ‘nunatak 1390’ comprises rocks with some of the best preserved primary textures and structures found in the tasiusarsuaq terrane. in brief, ‘nunatak 1390’ contains a volcanic series with variably preserved pillow lava sequences that are succeeded by melanocratic-ultramafic ash and rocks with flow structures (stendal & scherstén 2007). rhyolitic rocks 74 age (ga) 3.03.25 2.52.75 metamorphism, amphibolite/granulite grade emplacement of granite qôrqut granite complex emplacement of dioritic and ttg gneiss terrane accretion deposition of supracrustal rocks tre brødre isukasia færingehavn tasiusarsuaq akia kapisilik supracrustal units22 3311 2 31 fig. 2. time lines for igneous and metamorphic events in the nuuk region. numbers for the supracrustal units refer to (1) ivisaartoq and qussuk, (2) ‘nunatak 1390’ and (3) storø. data from pidgeon & kals beek (1978), friend & nutman (2001), crowley (2002), hanmer et al. (2002), friend & nutman (2005), polat et al. (2008), nutman & friend (2007), garde (2007) and knudsen et al. (2007). a b c d e f fig. 3. backscattered electron images of representative zircon grains for each dated sample. zircons were separated, picked, mounted in epoxy and polished to expose the central part of the grains. scale bars = 50 µm. (a) 499221 palaeosome; (b) 499221 neosome; (c) 499161; (d) 514832; (e) 499227; (f) 499228. are intercalated with the pillow lava sequences, and were interpreted as ignimbrites by stendal & scherstén (2007). in 2007, however, intrusive discordant dykes that appear to feed into the rhyolites were discovered, and it may be that some or all of these rocks are significantly younger sills. zircons were extracted from a rhyolite sample (499161) for dating, and in spite of the alternative interpretations either date the pillow lava sequence or provide a minimum age for it. the zircon grains have bright homogeneous cores and darker rims and/or zones within the cores (fig. 3c). the dark areas may extend into the bright areas and contain minor amounts of ca and al, which could indicate destabilisation and partial zircon breakdown. of 48 age determinations 43 define a recent pb loss line and a 207pb/206pb age of 2.873 ± 0.005 ga (n = 43/48, ± 2σ, mswd = 1.7). one grain (two analyses) is concordant at c. 3.2 ga. we interpret the 2.873 ±0.005 ga age to represent the time of crystallisation of the rhyolite. the origin of the 3.2 ga old grain is unclear, but we speculate that it is inherited from rocks of this age. gneisses with palaeosomes of this age have been found farther to the west in the tasiusarsuaq terrane (næraa & scherstén, unpublished data). eastern nunataarsuk is characterised by an anorthositeamphibolite-granite succession. granite sample 514832 for geochronology is from an area where granite and slices of amphibolite with local pillow structures in low-strain areas form the major successions; the granite-amphibolite contact is either an irregular and intrusive or a boudinaged contact parallel to the foliation, and the granite is interpreted as late to posttectonic relative to a locally defined s1 foliation (kolb & stendal 2007). zircon grains from the granite have oscillatory-zoned cores surrounded by thin homogeneous rims (fig. 3d). oscillatory-zoned grains plot along a recent pb-loss line and yielded a 207pb/206pb age of 2.852 ± 0.005 ga (n = 30/33, ± 2σ, mswd = 0.30, concordant within 10%), which we interpret as the intrusive age of the granite. the rims were generally too narrow to be analysed with our standard 20 µm laser spot; however, three analyses gave a poorly defined upper intercept age of 2.58 ± 0.04 ga. veined gneiss and cross-cutting tonalitic schist in the southern part of the region, a slightly schistose tonalite (499227) has an intrusive cross-cutting relationship to a veined tonalitic gneiss (499228) with amphibolite enclaves. abun dant veins of granitic pegmatite appear to have formed by partial melting of the veined gneiss along amphibolite bou din necks, along the foliation and along the tonalitegneiss contacts. the zircon grains from the tonalite (499227) have complex internal textures. many of them have oscillatory zoned cores and homogeneous rims and/or internal zones (fig. 3e). a total of 49 of 60 u-pb zircon age determi na tions define a 207pb/206pb age of 2.719 ± 0.005 ga (n = 49/60, ± 2σ, mswd = 2.2, concordant from 91 to 113%). the remaining 11 age determinations form an array towards older ages and presumably represent ancient pb-loss in grains as old as c. 2.859 ± 0.014 ga (single grain 207pb/206pb age; fig. 4e). the zircon grains from the veined gneiss (499228) generally have thick, homogeneous, bright rims surrounding cores that are highly cracked and commonly metamict; only a few cores show remains of oscillatory zonation (fig. 3f ). the spread in ages indicates ancient lead loss from 2.85 to 2.7 ga with no obvious 207pb/206pb age plateau (fig. 4f ); one grain was dated at c. 3.1 ga. the interpretation of the age data remains speculative, but given the altered appearance, the metamorphic rims and the relationship with the tonalite (499227), it seems reasonable to assume a metamorphic overprinting of the rock at c. 2.72 ga, perhaps related to the accretion with the tre brødre terrane. 75 499228 499221 neosome b 499227 3.00 20 7 p b/ 20 6 p b ag e (g a) 2.90 2.80 2.70 3.00 20 7 p b/ 20 6 p b ag e (g a) 20 7 p b/ 20 6 p b ag e (g a) 2.50 2.90 2.80 2.70 2.60 499221 paleosome a e f 514832499161 dc 2.90 2.66 2.82 2.86 2.78 2.74 2.70 fig. 4. zircon 207pb/206pb ages. all errors presented at 2 sigma levels. (a) sample 499221 palaeosome. the red bars represent data used to calculate crystallisation age (green horizontal bar). five blue bars are rejected due to assumed ancient lead loss. (b) sample 499221 neosome. the upper green horizontal bar represents the crystallisation age for the paleosome and the lower green horizontal bar represents the known granulite facies event for the tasiusarsuaq terrane. (c, d) samples 499161 and 514832. the red bars represent data used to calculate crystallisation ages (green horizontal bar) and the blue bars were rejected due to assumed ancient lead loss. (e) sample 499227: red bars represent data used to calculate crystallisation or metamorphic age (green horizontal bar), and the blue bars are assumed to be inherited from an older source. (f) sample 499228. the red bars define an ancient lead loss trend with age components as in sample 499227. age determinations were carried out on the element 2 laser icpms at geus. detailed analytical procedures are described in frei et al. (2006). discussion and summary the zircon crystallisation age of 2.87–2.85 ga for samples 499221, 499161 and 514832 and a somewhat speculative metamorphic age at 2.80 ga for sample 499221 correlate very well with known crystallisation ages and granulite facies events within the tasiusarsuaq terrane. these ages are significantly younger than ages of rocks from the kapisilik terrane but older than those of the tre brødre terrane. it is thus tempting to ascribe these areas to the tasiusarsuaq terrane. to include the ‘nunatak 1390’ is straightforward, while including nunataarsuk implies a major northerly extension of the terrane that requires confirmation by further work. age data alone naturally do not justify the inclusion of these areas into the tasiusarsuaq terrane, but it appears to be the most straightforward option based on the available information. more importantly, the ages obtained from the rhyolite and granite at ‘nunatak 1390’ (2.873 ga) and nunataarsuk (2.853 ga) provide minimum extrusive ages for the associated mafic greenstones and might reflect the onset of crustal growth in this block. furthermore, the emerging terrane configuration might indicate that the tasiusarsuaq terrane accreted with the kapisilik terrane in the north-east and with the tre brødre terrane in the south-west. the ages from the schistose tonalite (499227) and the veined gneiss (499228) are too young to readily represent known events within the tasiusarsuaq terrane. the inferred metamorphic ages rather correlate with the thermal event associated with prograde amphibolite facies metamorphism within the tre brødre terrane. however, situated well within the tasiusarsuaq terrane these rocks would not have experienced a prograde metamorphic path during terrane accretion. furthermore, there is evidence for ≥2.85 ga old zircons, which are too old to readily fit with known ages of the tre brødre terrane. we speculate that the northern part of the tasiusarsuaq terrane may represent a nappe complex, and that the investigated rocks either represent a tectonic window exposing footwall rocks that experienced prograde metamorphism and partial melting during overthrusting, or that fluids released from footwall-induced zircon pb-loss and/or partial melting in the overriding nappe. references coney, p.j., jones, d.l. & monger, j.w. 1980: cordilleran suspect terranes. nature 288, 329–332. crowley, j.l. 2002: testing the model of late archean terrane accretion in southern west greenland: a comparison of the timing of geological events across the qarliit nunaat fault, buksefjorden region. precambrian research 116, 57–79. escher, j.c. & pulvertaft, t.c.r. 1995: geological map of greenland, 1:2 500 000, copenhagen: geological survey of greenland. frei, d., hollis, j.a., gerdes, a., harlov, d., karlsson, c., vasquez, p., franz, g., johansson, l. & knudsen, c. 2006: advanced in situ geochronological and trace element microanalyses by laser ablation techniques. geological survey of denmark and greenland bulletin 10, 25–28. friend, c.r.l. & nutman, a.p. 2001: u-pb zircon study of tectonically bounded blocks of 2940–2840 ma crust with different metamorphic histories, paamiut region, south-west greenland: implications for the tectonic assembly of the north atlantic craton. precambrian research 105, 143–164. friend, c.r.l. & nutman, a.p. 2005: new pieces to the archaean terrane jigsaw puzzle in the nuuk region, southern west greenland: steps in transforming a simple insight into a complex regional tectonothermal model. journal of the geological society (london) 162, 147–162. friend, c.r.l., nutman, a.p., baadsgaard, h., kinny, p.d. & mcgregor, v.r. 1996: timing of late archaean terrane assembly in the nuuk region, southern west greenland. earth and planetary science letters 142, 353–365. garde a.a. 2007: a mid-archaean island arc complex in the eastern akia terrane, godthåbsfjord, southern west greenland. journal of the geological society (london) 164, 565–579. knudsen, c., van gool, j.a.m., østergaard, c., hollis, j.a., rink-jørgensen, m., persson, m. & szilas, k. 2007: gold-hosting supracrustal rocks on storø, southern west greenland: lithologies and geological environment. geological survey of denmark and greenland bulletin 13, 41–44. kolb, j. & stendal h. 2007: geological environments and hydrothermal mineralisation in nunataarsuk, qarliit nunaat and ameralik, nuuk region, sw greenland – a field report 2007. mineral resource assessment of the archaean craton (66° to 63°30´n), sw greenland. contribution no. 2. danmarks og grønlands geologiske undersøgelse rapport 2007/58, 44 pp. nutman, a.p. & friend, c.r.l. 2007: adjacent terranes with ca. 2715 and 2650 ma high-pressure metamorphic assemblages in the nuuk region of the north atlantic craton, southern west greenland: complexities of neoarchaean collisional orogeny. precambrian research 155, 159–203. nutman, a.p., friend, c.r.l., baadsgaard, h., & mcgregor, v.r. 1989: evolution and assembly of archaean gneiss terranes in the godt håbs fjord region, southern west greenland: structural, metamorphic and isotopic evidence. tectonics 8, 573–589. pidgeon, r.t. & kalsbeek, f. 1978: dating of igneous and metamorphic events in the fiskenaesset region of southern west greenland. canadian journal of earth sciences 15, 2021–2025. polat a., frei, r., appel, p.w.u., dilek, y., fryer, b., ordóñez-calderón, j.c. & yang, z. 2008: the origin and composition of mesoarchean oceanic crust: evidence from the 3075 ma ivisaartoq greenstone belt, sw greenland. lithos 100, 293–321. stendal, h. & scherstén, a. 2007: a well-preserved bimodal archaean volcanic succession in the tasiusarsuaq terrane, south-west greenland. geological survey of denmark and greenland bulletin 13, 53–56. 76 authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: tomn@geus.dk geological survey of denmark and greenland bulletin 15, 2008, 9-12 upper jurassic quartz-rich sandstones in the north sea basin are important reservoir rocks for oil and gas, and one of the latest discoveries of oil in the danish sector was made in the area of the hejre wells that penetrated such sediments (fig. 1). the reservoir properties of sandstones are strongly influenced by diagenetic alteration, i.e. the mineralogical changes that take place during burial of the sediments. the diagenetic features depend on the source area, depositional setting, facies architecture and burial history of the sediment. the major diagenetic features influencing porosity in upper jurassic reservoir sandstones are feldspar dissolution and precipitation, preci-pitation of illite, calcite and quartz, and quartz stylolite formation. with regard to the upper jurassic sandstones in the danish sector of the north sea, the important question is: how can porosity be preserved in sediments buried at depths of more than 5 km? the hejre-2 well penetrated the upper jurassic sediments (fig. 2) before reaching pre-upper jurassic volcaniclastic conglomerates. the diagenetic features were studied in thin sections of core samples with traditional petrographic techniques using transmitted light microscopy supplemented by scanning electron microscopy (sem) of rock chips and thin sections. feldspar dissolution and precipitation previous studies of feldspar diagenesis have generally focused on the albitisation of k-feldspar, the dissolution and replacement of plagioclase with clay at deep burial depths, and their importance for the generation of secondary porosity (e.g. surdam et al. 1984; bjørlykke et al. 1992). recent investigations show that the formation of feldspar overgrowths may also be a significant early diagenetic phenomenon (lee & parsons 2003). early diagenetic authigenic k-feldspar has also played an important role in the upper jurassic sandstones investigated here, although the volume of precipitated material is not nearly as high as the succeeding calcite and quartz cementation. detrital feldspar, and notably k-feldspar, is a common constituent of upper jurassic sediments. authigenic kfeldspar may have formed during early diagenesis, but is now mainly preserved in calcite-cemented parts of the sandstones, where it forms overgrowths on detrital k-feldspar, perthite and albite grains (fig. 3a, b). authigenic k-feldspar is com© geus, 2008. geological survey of denmark and greenland bulletin 15, 9–12. available at: www.geus.dk/publications/bull diagenesis influencing the porosity of upper jurassic reservoir sandstones, danish north sea rikke weibel and nynke keulen 9 0° 0° 8° 58°n58°n 50°n 8°e germanyunited kingdom north sea denmark hejre wells danish central graben 500 km fig. 1. sketch map of the north sea region showing the location of the hejre wells. the dashed lines show the borders between the national sectors. sy st em se ri es ju r a ss ic u pp er m id dl e stage kimmeridgian oxfordian callovian bathonian central graben british sector danish sector norwegian sector fu lm ar f m fu lm ar f m farsund fm farsund fm u la f m haugesund fm lola fm bryne fm bryne fm middle graben fm gert mb ravn mb h en o fm k im m er id ig e c la y fm lulu fm r at tr ay v ol ca ni c m b pe nt la nd f m r on v ol ca ni c m b sand-rich sediments clay-rich sediments hiatus fig. 2. middle and upper jurassic lithostratigraphy of the british, danish and norwegian sectors of the central graben in the north sea. the hejre-2 well penetrated the farsund formation, the ravn member of the heno formation, the lola formation and the gert member of the heno formation. simplified after johannesen et al. (2003) and michelsen et al. (2003). mon in the vicinity of k-bearing clay-rich intervals and in the sandstone immediately overlying volcaniclastic conglomerates. element mapping of the calcite-cemented parts of the sandstones shows the volumetric importance of the early authigenic k-feldspar phase (fig. 3b). the marine upper jurassic sediments (johannessen 2003) are expected to have had relatively high concentrations of potassium in the original pore fluids, in the k-feldspar stability field (e.g. worden & morad 2000). however, the large proportions of authigenic k-feldspar associated with clay-rich intervals and the underlying volcaniclastic conglomerates could indicate an additional external supply of potassium. the volcaniclastic conglomerates have a very high k-content and must originate from a potassic alkaline volcanic source. lee & parsons (2003) found exceptionally high proportions of sanidine crypto perthites in the upper jurassic humber group, inferred to have an ultra-potassic volcanic source. potassic volcanic sources of early and middle jurassic (furnes et al. 1982) or permian age (aghabawa 1993) are present in the north sea area, but a determination of the age and location of the actual volcanic source of potassium for the upper jurassic reservoir sandstones require further investigations. remnants of albite grains are only preserved in the calcitecemented parts of the sandstones, whereas more calcium-rich plagioclase detrital grains are almost completely lacking in the sandstones. these parts of the sandstone are instead characterised by k-feldspar overgrowths around voids of approximately the size of average detrital grains (fig. 3c). these voids are often partly filled with illite and occasionally authigenic quartz. the k-feldspar overgrowths presumably precipi10 25 µm 100 µm a b c d 100 µm 20 µm qtz-o oil oil illite + oil qtz-o kfsp kfsp i i i kfsp fig. 3. images showing diagenetic alteration of feldspars in upper jurassic sandstones that resulted in both reduction and increase of porosity. a: abundant detrital k-feldspar and k-feldspar overgrowths (arrows) on detrital albite in calcite-cemented parts of the sandstones. back-scattered electron sem image. b: combined element maps of the area shown in fig. 3a. orange, calcium; yellow, potassium; pink, sodium; blue, quartz. c: k-feldspar overgrowths (kfsp) most likely formed on an original plagioclase grain prior to its dissolution and replacement by illite (i) during which substantial secondary porosity was created. subsequent quartz overgrowths (qtz-o) and compression may have reduced this porosity. secondary electron sem image. d: detrital k-feldspar (kfsp) partly dissolved (white arrows) and replaced by illite (i). some oil droplets (oil) can be observed. secondary electron sem image. tated around detrital albite or ca-rich plagioclase grains that were subsequently dissolved or replaced by clay minerals, mainly illite. it appears that dissolution of detrital albite and ca-rich plagioclase grains led to substantial secondary porosity in the non-calcite cemented parts of the sandstones; even kfeldspar shows partial dissolution and illitisation (fig. 3d). however, subsequent compaction and growth of authigenic phases have reduced this secondary porosity (fig. 3b). further more, it is possible that the relatively high porosity in the lower parts of the sandstones originates from early precipitation of authigenic k-feldspar cement, which, similar to the detrital feldspars, was subsequently dissolved or replaced by illite re sulting in major secondary porosity. quartz diagenesis: stylolites, pressure solution and overgrowth quartz cementation by development of stylolites, pressure solution between adjacent quartz grains and precipitation of macro-overgrowths are important in deeper (3–5 km) buried siliciclastic reservoirs (e.g. bjørlykke et al. 1992; worden & morad 2000). the investigated upper jurassic sediments are no exception, as detrital quartz grains are the major constituent of the upper jurassic sediments and quartz diagenesis is the major porosity-reducing feature outside the calcitecemented areas. however, parts of the sandstones show a high porosity and are without major quartz or calcite cement. the porosity is reduced by the formation of stylolites (fig. 4a), pressure solution and by precipitation of quartz overgrowths (fig. 4b, c); all late diagenetic processes that took place after authigenic k-feldspar precipitation, calcite cementation and feldspar dissolution and illitisation. the stylolites appear to have evolved from primary mica-rich and detrital clay-rich (mainly illitic) laminae (fig. 4a). quartz pressure solution and stylolite formation may be related to the local microphysiochemical environment where alkaline conditions formed around clays, which could lead to enhanced pressure solution of quartz through increased solubility or an increased dissolution rate (tada & siever 1989). pressure solution has also been related to the surface charge of the illitic clays or mica (walderhaug et al. 2006). illitic clays are quite common in the investigated upper jurassic sandstones and may have led to enhanced development of stylolites. on the other hand, in sediments with high concentrations of clay and rock fragments, stylolites may be completely absent, due to the cushioning effect of the clay matrix and ductile lithic fragments (cf. tada & siever 1989). the pressure solution of detrital quartz grains and stylolite formation may have released silica for the quartz overgrowths. syntaxial quartz overgrowths are found in all parts of the upper jurassic sediments except for the calcite-cemented areas. peculiar, irregularly shaped quartz overgrowths are found in those parts of the sediments that show the highest present-day porosities (fig. 4b, c). these peculiar quartz overgrowths 11 500 µm 50 µm 50 µm stylolite qtz-o qtz-o qtz-o qtz-o clay clay qtz-o stylolite qtz-o qtz-o qtz-o qtz-o clay clay qtz-o a b c fig. 4. images showing development of quartz stylolites, quartz pressure solution and quartz overgrowths that are the main diagenetic features affecting the porosity outside the calcite-cemented parts of the sandstones. a: quartz stylolite formation is enhanced by the presence of detrital clays and mica. non-calcite cemented parts of the sandstones. trans mitted light microscope image. b: quartz pressure solution between adjacent detrital quartz grains (black arrows) and quartz overgrowths (qtz-o) that locally appear to be partly inhibited due to thick coatings of authigenic illite (white arrows). non-calcite cemented sandstone. back-scattered electron sem image. c: same area as fig. 4b, but cathodoluminescence sem image. can be interpreted as an indication of three completely different processes. they could either be caused by an incipient introduction of hydrocarbons, which partly covered the overgrowth (cf. marchand et al. 2000). in other cases, it has been suggested that authigenic chlorite rims have inhibited or retarded quartz overgrowth (aase et al. 1996). aase & walderhaugh (2005) argued that the delay of such quartz precipitation was caused by thin coatings of micron-scale quartz crystals (microquartz). in sandstones of the hejre-2 well abundant illite, formed during alteration of feldspar, may have inhibited pressure solution and stylolite formation, as a result of the cushioning effect of the clays, and furthermore have delayed the growth of authigenic quartz in the upper jurassic sediments. summary and perspectives the quartz-rich upper jurassic reservoir sandstones in the hejre-2 well are characterised by relatively high amounts of detrital k-feldspar, which most likely come from a volcanic source. authigenic k-feldspar grew on detrital k-feldspar, perthite and albite grains and possibly ca-rich plagioclase. the potassium for the k-feldspar overgrowths probably comes from marine pore fluids, detrital clay minerals and the underlying volcaniclastic conglomerates. the k-feldspar over growths were initially abundant, but were later dissolved except for the calcite-cemented parts of the sandstones. the formation of secondary porosity in the sandstones in the hejre-2 well was mainly caused by dissolution of detrital feldspars and authigenic k-feldspar. quartz diagenesis prevails in the upper jurassic sediments in the hejre-2 well where carbonates are absent and where illitic clays are restricted. pressure solution and stylolite formation provide abundant silica, which is precipitated as quartz overgrowths. the presence of small amounts of illitic clays and mica may have enhanced the development of stylolites. on the other hand, relatively large amounts of illitic coatings seem to have inhibited or retarded the formation of quartz overgrowths, which otherwise could have reduced the porosity. consequently, sandstone intervals in the hejre-2 well with high porosities are the result of dissolution and illite replacement of detrital feldspars and early authigenic k-feld spar cement, thus secondary porosity was created; together with illite coatings that inhibited the formation of quartz overgrowths. future investigations may reveal whether the diagenetic evolution of upper jurassic sediments in other wells is similar to that encountered in the hejre-2 well. another aspect worth pursuing is to investigate if bulk rock geochemistry can be correlated with major diagenetic phenomena. acknowledgement dong energy is thanked for permission to publish the present information on the hejre-2 well prior to public release of the data. references aase, n.e., bjørkum, p.a. & nadeau, p.h. 1996: the effect of grain-coating microquartz on preservation of reservoir porosity. american association of petroleum geologists bulletin 80, 1654–1673. aase, n.e. & walderhaug, o. 2005: the effect of hydrocarbons on quartz cementation: diagenesis in the upper jurassic sandstone of the miller fields, north sea, revisited. petroleum geoscience 11, 215–223. aghabawa, m.a. 1993: petrology and geochemistry of the rotliegendes volcanic rocks in denmark and their tectonic implications. dyn maisk/stratigrafisk analyse af palæozoikum i danmark. efp-89; område 1: olie og naturgas, 3, dgu kunderapport 1993/35, 351 pp. bjørlykke, k., nedkvitne, t., ramm, m. & saigal, g.c. 1992: diagenetic processes in the brent group (middle jurassic) reservoirs of the north sea: an overview. in: morton, a.c. et al. (eds): geology of the brent group. geological society special publication (london) 61, 263–287. furnes, h., elvsborg, a. & malm, o.a. 1982: lower and middle jurassic alkaline magmatism in the egersund sub-basin, north sea. marine geology 46, 53–69. johannessen, p.n. 2003: sedimentology and sequence stratigraphy of paralic and shallow marine upper jurassic sandstones in the northern danish central graben. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 367–402. lee, m.r. & parsons, i. 2003: microtextures of authigenic or-rich feldspar in the upper jurassic humber group, uk north sea. sedimentology 50, 597–608. marchand, a.m.e., haszeldine, r.s., macaulay, c.i., swennen, r. & fallick, a.e. 2000: quartz cementation inhibited by crestal oil charge: miller deep water sandstone, uk north sea. clay minerals 35, 201–210. michelsen, o., nielsen, l.h., johannessen, p.n. andsbjerg, j. & surlyk, f. 2003: jurassic lithostratigraphy and stratigrafic development onshore and offshore denmark. in: ineson, j.r. & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 147–216. surdam, r.c., boese, s.w. & crossey, l.j. 1984: the chemistry of secondary porosity. in: mcdonald, d.a. & surdam, r.c. (eds): clastic diagenesis. american association of petroleum geologists memoir 37, 127–149. tada, r. & siever, r. 1989: pressure solution during diagenesis. annual review of earth and planetary science 17, 89–118. walderhaug, o., bjørkum, p.a. & aase, n.e. 2006: kaolin-coating of stylolites, effect on quartz cementation and general implications for dissolutions at mineral interfaces. journal of sedimentary research 76, 234–243. worden, r.h. & morad, s. 2000: quartz cementation in oil field sandstones: a review of the key controversies. in: worden, r.h. & morad, s. (eds): quartz cementation in sandstones. special publication of the international association of sedimentologists 29, 1–20. oxford: blackwell science. authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: rwh@geus.dk 12 piasecki 2024: geus bulletin 57. 8373. https://doi.org/10.34194/geusb.v57.8373 1 of 15 research article the scriniodinium crystallinum dinoflagellate cyst zone in the middle–upper oxfordian, upper jurassic, ilimananngip nunaa (milne land), east greenland stefan piasecki* emeritus, section for geobiology, globe institute, university of copenhagen, copenhagen, denmark; department for geophysics and sedimentary basins, geological survey of denmark and greenland (geus), copenhagen, denmark abstract the biostratigraphy of the jurassic in east greenland is historically based on macroscopic fossils. stratigraphy based on palynomorphs (spores, pollen and dinoflagellate cysts) has progressed more slowly and sporadically. the scriniodinium crystallinum dinoflagellate cyst zone is identified in middle – upper oxfordian strata of ilimananngip nunaa (milne land), central east greenland. the lower boundary is defined by the last occurrence of trichodinium scarburghense in the cardioceras tenuiserratum ammonite zone. the upper boundary is defined by the last occurrence of s. crystallinum in the uppermost amoeboceras rosenkrantzi ammonite zone. however, the subzonal division of the s. crystallinum zone recorded in north-west europe is not identified in greenland. eighteen characteristic dinoflagellate cyst events are considered stratigraphically significant and useful in east greenland. fifteen of these events provide an informal, detailed stratigraphical subdivision of the s. crystallinum zone into 10 subunits. identification of the zone is an addition to the previously defined upper bathonian – middle oxfordian zonation, where the uppermost palynostratigraphical event was recorded to be the last occurrence of t. scarburghense. with this study, the correlation of dinoflagellate cyst and ammonite stratigraphy in the lower and middle oxfordian is slightly modified. the s. crystallinum zone documented here, in combination with the zonation used for the stratigraphy of the blokelv-1, rødryggen-1 and brorson halvø-1 cores of the upper jurassic to lower cretaceous, completes the dinoflagellate cyst stratigraphy of the marine jurassic in east greenland. together with previous studies of spores and pollen in less marine units, the first complete palynological jurassic stratigraphy is thus established for the jurassic succession in east greenland. *correspondence: stefan.piasecki@sund. ku.dk received: 23 feb 2024 revised: 06 may 2024 accepted: 09 jun 2024 published: 04 nov 2024 keywords: dinoflagellate cysts, stratigraphy, upper jurassic, oxfordian– kimmeridgian, east greenland geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: karen dybkjær (geus, denmark) reviewers: james b. riding (british geological survey, uk), morten smelror (geological survey of norway) funding: see page 14 competing interests: see page 14 additional files: none 1 introduction jurassic biostratigraphy is historically based mainly on macroscopic fossils, especially on ammonites but with significant contributions from other fossil groups. the jurassic ammonite stratigraphy became the standard stratigraphical zonation, and the ammonite zones were applied as chronozones (e.g. callomon 1984, 1993). integration and correlation of microscopic fossil stratigraphy with ammonite stratigraphy developed and accelerated in the 1950s–1960s, following increasing demand of stratigraphical frameworks from expanding, worldwide industrial drilling programs primarily for offshore energy exploration. studies of dinoflagellate cysts and their stratigraphical occurrences were established relatively early for the jurassic in europe but with somewhat slower progress in the northern atlantic region. on jameson land, east greenland (fig. 1), jurassic sedimentary samples for dinoflagellate cyst stratigraphy were collected systematically from beds with ammonites by tove birkelund in the 1970s and made accessible to students and industry palynologists. on ilimananngip nunaa (milne land) in 1977, samples for palynological studies were collected mainly by this author, with contributions from members https://doi.org/10.34194/geusb.v57.8373 https://orcid.org/0000-0002-7846-859x mailto:stefan.piasecki@sund.ku.dk mailto:stefan.piasecki@sund.ku.dk https://creativecommons.org/licenses/by/4.0/deed.ast piasecki 2024: geus bulletin 57. 8373. https://doi.org/10.34194/geusb.v57.8373 2 of 15 geusbulletin.org of the field team: t. birkelund, john h. callomon (ammonites), claus heinberg and franz fürsich (molluscs) and lars stemmerik (sedimentology). successions at 52 localities were sedimentologically logged, collected for fossils and sampled for palynology, coordinated by three field teams. the studied localities were numbered 1–52 (piasecki 1980, fig. 1; birkelund & callomon 1985, fig. 1). locality 39 east of visdal on milne land (fig. 1c) was measured in 1977, and a series of samples were marked by birkelund & callomon (1985, figs 1 and 3) and collected (piasecki 1980, fig. 34). the succession belongs to the kosmocerasdal, aldinger elv and bays elv members of the kap leslie formation (fig. 2). five faunal horizons of ammonites were identified and referred to the milne land ammonite faunal horizons m 12, m 13 and m 14 (fig. 3; callomon & birkelund 1980; birkelund & callomon 1985). these faunas represent the boreal faunal province (amoeboceras regulare and amoeboceras rosenkrantzi ammonite zones) and the sub-boreal province (pictonia baylei ammonite zone) of the upper oxfordian to the lowermost kimmeridgian (fig. 3). such a complete ammonite zonation has not been documented in any other exposure on milne land or jameson land, though faunal horizons m 13 and m 14 were tentatively identified in the fine-grained hareelv formation on jameson land (callomon & birkelund 1980, fig. 3). faunal horizons are not recorded in the lower part of the succession at locality 39 of kosmocerasdal member (fig. 2). comparison with other exposures in eastern milne land indicates that this part of kosmocerasdal member presumably correlates with the upper oxfordian, uppermost amoeboceras glosense – amoeboceras serratum ammonite zones. faunal horizon m 11, a. serratum ammonite zone, is only recognised in coarse-grained, sandstone exposures of aldinger elv member (fürsich & heinberg 1983), where no samples were collected for dinoflagellate cysts. locality 39 displays a condensed and continuous upper oxfordian succession, well dated by ammonites, and is probably the only complete section in central east greenland. the sediment is mostly fine-grained sandstone with concretions and little potential for palynological content. nevertheless, dinoflagellate cyst assemblages have been recovered from all samples. for this study, palynological sample materials from ammonite bearing beds and horizons were collected for direct comparison of palynology and ammonite stratigraphy in the oxfordian from a composite section from kosmocerasdal (locality 2), nordøstelv (locality 3), ‘ilovaiskii’ dal (locality 4) and ‘hystrix’ dal (locality 5) – all below aldinger elv member in the north-eastern exposures of jurassic sediments on milne land (figs 1 and 4). this composite succession is integrated with that of locality 39 and extends the jurassic succession in east greenland downwards into the existing dinoflagellate cyst stratigraphy of smelror (1988). 2 material and methods all material was originally sampled in the field campaign of 1977 from the kap leslie formation at localities 2–6 and 39 (locations in fig. 1). the kap leslie formation is dominated by conglomerates, sand and muddy sandstone. only the gråkløft member of the kap leslie formation comprises laminated dark mudstones (figs 2 and 4). the stratigraphical range chart presented here (fig. 5) is a combination of two successions. the lower part is from localities 2–5 (shown at reduced scale in fig. 5 and true scale in fig. 4). the upper part is from locality 39 (illustrated at true scale in fig. 5). the bays elv member above the aldinger elv member in cardioceraskløft (locality 6) has no confidently identified ammonite faunal horizons but correlates with the upper succession of locality 39 (fig. 4; piasecki 1980; birkelund & callomon 1985). stratigraphical events from the cardioceraskløft succession (fig. 4) are compared here with data from locality 39, to support the new palynological zone for the upper oxfordian but are not included in the range chart in fig. 5. the ammonites were generally recovered in carbonate cemented sandstone beds or calcareous concretions and were subsequently referred to ammonite faunal horizons, m 1–m 15, by callomon & birkelund (1980; e.g. m 2 is the second faunal horizon on milne land). the faunal horizons are referred to ammonite zones (fig. 3). the ammonite stratigraphy applied here follows the east greenland tradition (e.g. sykes & surlyk 1976; callomon 1993) of combined boreal and sub-boreal zonation (fig. 3) and likewise the concept of considering ammonite zones as chronozones. here, however, ammonite zones are consequently referred to as biozones. the basis of an ammonite zone is indicated at the lowermost occurrence of the index species in a faunal horizon, and the top of the zone is defined by the appearance of the next index species. compared to earlier interpretations, the cardioceras densiplicatum and cardioceras tenuiserratum ammonite zones are consequently reduced slightly in thickness and sample density, whereas the thickness of a. glosense ammonite zone is expanded. the sample material was processed by standard methods. the crushed samples were prepared with acid (hcl, hf and hno3) to remove carbonate and silica (clay, silt and sand) from the samples. organic matter is resistant to the acid, and the remains contained abundant terrestrial organic material, especially brown and black woody material. the separation method developed by hansen & gudmundsson (1979) was applied to the organic residue and successfully removed most of the abundant woody material and improved the recovery of identifiable dinoflagellate cysts significantly. the remaining organic residue was mounted in glycerine-gelatine on preparation glasses for visual analysis using a standard light microscope. although the slides were prepared in the 1970s, most are https://doi.org/10.34194/geusb.v57.8373 http://www.geusbulletin.org/ piasecki 2024: geus bulletin 57. 8373. https://doi.org/10.34194/geusb.v57.8373 3 of 15 geusbulletin.org volquart boon coast crystalline basement kap leslie fm charcot bugt fmhartz fjeld fm paleogene basalt hesteelv fm raukelv fm hareelv fm pelion fm neill klinter gp kap stewart gp 23°w24°w 70°30ʹn 25 km jameson land b blokelv-1 18°w 16°w 76°n 72°n 70°n 22°w 20°w saf pdmf df lle lle cretaceous jurassic triassic permian main faults stauning alper fault post-devonian main fault dombjerg fault liverpool land escarpment locality greenland ilim ananngip nunaa (m ilne land) gåseland traill ø geographical society ø hold with hope clavering ø th.thomsen land kuhn ø hochstetter forland store koldewey germania land a kulhøj wollaston forland p d m f s a f d f jameson land kangerluk kong oscar liverpool land kangersaajiva (hurry inlet) 50 km 26°w 22°w24°w28°w 26°w 24°w hall bredning 74°n bc a brorson halvø -1 rødryggen-1 25°20ʹw 70°50ʹn 20 km 25°40ʹw c 2 3 4 5 6 39 ! kangertittivaq (scoresby sund) ilimananngip nunaa (milne land) k an ge rs aa jiv a (h ur ry in le t) quaternary ice sea rivers faults fig. 1 map of the study area. a: geological map of east greenland (modified from surlyk et al. 2021) with locations of the rødryggen-1 and brorson halvø-1 drilling sites. inset: maps of localities analysed in this study. b: jameson land with blokelv-1 core and c: ilimananngip nunaa (milne land) with localities m 2, 3, 4, 5, 6 and 39 (locality numbers sensu birkelund et al. 1984). https://doi.org/10.34194/geusb.v57.8373 http://www.geusbulletin.org/ piasecki 2024: geus bulletin 57. 8373. https://doi.org/10.34194/geusb.v57.8373 4 of 15 geusbulletin.org still perfectly preserved, and only a few had to be restored for the present study. the dinoflagellate cyst taxonomy follows fensome et al. (2019) and riding et al. (2022) with respect to the former gonyaulacysta jurassica group. 2.1 history of dinoflagellate cyst stratigraphy in east greenland overall, the middle – upper jurassic, boreal or sub-boreal ammonite stratigraphy in east greenland is well documented on milne land (callomon & birkelund 1980; birkelund et al. 1984; birkelund & callomon 1985), jameson land (surlyk et al. 1973; callomon et al. 2015) and in north-east greenland (sykes & surlyk 1976). the studies mentioned in the following focus on middle – upper jurassic dinoflagellate cysts from central east greenland. the number of studies is limited, and only a few cover the middle – upper oxfordian interval. samples collected from central jameson land by r.c. whatley and d.c. brown in 1964 were prepared in the laboratory of university of nottingham, uk, but were subject to a fire in the laboratory. only parts of two preparations survived and were analysed by w.a.s. sarjeant (1972). but the value of these data suffers from uncertainties of their stratigraphical derivation. the age of the two samples was considered bathonian to callovian of the vardekløft formation, now vardekløft group (surlyk et al. 2021). the samples were collected near the locality langryggen in central jameson land, probably from within the lowermost olympen formation (athene and hades members) to fossilbjerget formation interval (see larsen & surlyk 2003; surlyk et al. 2021). the uppermost sample comprises endoscrinium luridum, trichodinium scarburghense and wanaea digitata and several other species that are in accordance with a latest callovian – earliest oxfordian age of the quenstedtoceras lamberti – quenstedtoceras mariae ammonite zones. the age of the dinoflagellate cyst assemblage of the lower sample (c. 75 m lower in the section) is imprecise. the presence of gonyaulacysta eisenackii, rhynchodiniopsis (gonyaulacysta) cladophora and pareodinia prolongata as well as a more marine assemblage may indicate that the sample is from the uppermost fossilbjerget formation of callovian age (see surlyk et al. 2021). a new acritarch genus (later referred to dinoflagellate cysts) mendicodinium (thuledinium) groenlandicum was described by pocock & sarjeant (1972) based on type material from sarjeant’s upper sample. the species is abundant in this sample, and it is known to have peak abundance in the q. lamberti ammonite zone (e.g. smelror 1988), indicating that the sample is from this ammonite zone, and consequently that both of sarjeant’s samples are of callovian age. a geographically and stratigraphically wide suite of 27 jurassic samples collected by t. birkelund on jameson lithostratigraphy of jurassic to lower cretaceous on milne land k ap l es lie f or m at io n kosmocerasdal member aldinger elv member bays elv member cardioceraskløft member gråkløft member krebsedal member pernaryggen member astartedal member hartz fjeld formation coarse-grained sand and conglomerates sand silty, very fine-grained sand pinnadal formation charcot bugt formationmiddle oxfordian to lower bathonian middle volgian to middle callovian valanginian to middle volgian hauterivian laminated, organic-rich mudstone chronostratigraphy fig. 2 lithostratigraphical scheme of jurassic to lower cretaceous sediments on milne land based on callomon & birkelund (1980), birkelund & callomon (1985), birkelund et al. (1984) and surlyk et al. (2021). kimmeridgian oxfordian callovian lo w er m id dl e u pp er m 14 m 15 m 13 m 12 m 11 m 10 m 9 m 8 m 7 m 6 m 5 m 4 m 3 m 2 stage fauna horizon ammonite zone ammonite fauna horizons, milne land borealsub-boreal ammonite faunal provinces: rasenia cymodoce pictonia baylei amoeboceras rosenkrantzi amoeboceras regulare amoeboceras serratum amoeboceras glosense amoeboceras glosense cardioceras tenuiserratum cardioceras densiplicatum cardioceras densiplicatum cardioceras densiplicatum cardioceras cordatum quenstedtoceras mariae quenstedtoceras lamberti peltoceras athleta fig. 3 ammonite faunal horizons on milne land correlated with ammonite zones (callomon & birkelund 1980; birkelund & callomon 1985). shading indicates the faunal province affiliations. https://doi.org/10.34194/geusb.v57.8373 http://www.geusbulletin.org/ piasecki 2024: geus bulletin 57. 8373. https://doi.org/10.34194/geusb.v57.8373 5 of 15 geusbulletin.org m 3 m 4 m 2 m 2 m 5 m 6 m 8 m 8 m 9 m 9 m 10 m 10 locality 2 kosmocerasdal locality 3 nordøstelv locality 4 “ilovaiskii” dal locality 5 “hystrix” dal c ha rc ot b ug t f m k ap l es lie f m / k os m oc er ad al m b k os m oc er ad al m b k os m oc er ad al m b k os m oc er ad al m b k os m oc er ad al m b a ld in ge rs e lv m b c la y s ilt s an d g ra ve l c la y s ilt s an d g ra ve l c la y s ilt s an d g ra ve l c la y s ilt s an d g ra ve l c la y s ilt s an d g ra ve l 240 230 220 210 200 190 240 250 260 270 280 290 50 40 30 20 10 0 20 30 40 50 60 20 10 – 245977-78 – 245975-76 – 245974 – 245973 – 245972 – 245969-71 – 245968 – 245967 – 245966 – 245965 – 245964 – 245963 – 245962 – 245961 – 245958-60 – 245957 – 245956 – 245955 – 245954 – 245953 – 245952 – 245951 – 245990 – 245989 – 245987-88 – 245986 – 245985 – 245984 – 245983 – 245982 – 245981 – 245980 – 234242 – 234241 – 234240 – 234239 – 234238 – 234237 – 234235 – 234236 – 234172 – 234171 – 234169-70 – 234168 – 234167 – 234166 – 234165 – 234164 – 234163 – 234162 – 234161 – 234160 – 234159 – 234158 – 234157 – 234156 – 234155 – 234154 – 234153 – 234178 – 234177 – 234176 – 234175 – 234173-74 [m] [m] [m] [m] [m] m 15 m 7 m 12 m 13 m 13 m 14 m 13 10 20 30 40 0 locality 6 cardioceraskløft locality 39 e of visdal legend grain size: concretions: pebbles gravel sand silt sediment structures: giant-scale planar cross-bedding large-scale planar cross-bedding large-scale trough cross-bedding small-scale ripple cross-bedding lenticular bedding planar lamination pyrite glauconite basalt lithology: trace fossils: fossils: doggers calcareous or pyritic concretions ferro-concretions concretionary beds of distinct horizons of concretions wood ammonite belemnite bivalve diplocraterion thallasinoides chondrites planolithes muensteria – 245834 – 245833 – 245832 – 245831 – 245830 – 245829 – 245828 – 245827 – 245826 – 245825 – 245824 – 245823 – 245822 – 245818 – 245817 – 245816 – 234219 – 234220 – 234221 – 234224 – 234225 – 234227 – 234228 – 234229 – 234230 – 234231 – 234232 c la y s ilt s an d g ra ve l c la y s ilt s an d g ra ve l 100 90 80 70 [m] [m] a .e . m b b ay s e lv m em be r c . m b c ha rc ot b ug t f m k ap l es lie f or m at io n v is da l m b k os m oc er as da l m b a .e . m b b ay s e lv m em be r fig. 4 measured and sampled sedimentary successions from milne land with ammonite faunal horizons (e.g. m 14) and palynological samples (e.g. ggu245951) modified from piasecki (1980). the locality numbers are from piasecki (1980, fig.3) and birkelund & callomon (1985, fig. 2). the sedimentary succession from localities 2–5 shows the complete kosmocerasdal member, kap leslie formation. sedimentary logs from locality 6, cardioceraskløft, and locality 39, east of visdal, are correlated by the upper boundary of aldinger elv member and the basis of bays elv member. https://doi.org/10.34194/geusb.v57.8373 http://www.geusbulletin.org/ piasecki 2024: geus bulletin 57. 8373. https://doi.org/10.34194/geusb.v57.8373 6 of 15 geusbulletin.org land were published by fensome (1979). this comprehensive, taxonomic and excellent pioneering work in east greenland included four samples presumed to be of upper oxfordian – lower kimmeridgian from southern and middle jameson land. these samples could be expected to overlap stratigraphically with the material from middle – upper oxfordian to lowermost kimmeridgian of the present study. however, the proposed age of three of these samples was not based on correlation with ammonite zonation but on their presumed lithostratigraphical affiliation to the hareelv formation. the presence of wanaea fimbriata, w. digitata, rigaudella aemula and t. scarburghense in these samples (fensome 1979) suggests an earliest oxfordian age. these samples are therefore stratigraphically comparable with the lower olympen formation, especially the dark shales of the hades member (see larsen & surlyk 2003; surlyk et al. 2021). two of the samples were collected close to sarjeant’s locality in central jameson land. the fourth sample was referred to the ‘?decipia ammonite zone, lower kimmeridgian, but this zone is classed as upper oxfordian in the stratigraphical scheme (fensome 1979, table 1 after surlyk et al. 1973). the sample is assigned to the middle – upper oxfordian based on the dinoflagellate cyst content, and the species are also common in the assemblages of the present study. distinction of lower oxfordian black shales of the olympen formation from the upper oxfordian hareelv formation is problematic in the south to mid-jameson land basin, where the shales are deposited successively. the two dark shale units cannot be distinguished lithologically from each other but only by biostratigraphy (finn surlyk, pers. comm. 2023). piasecki (1980) analysed dinoflagellate cysts from callovian to middle volgian on milne land. the correlation of dinoflagellate cyst stratigraphy with ammonite zones is revised slightly herein for a few of these samples (fig. 4). the results have been published in various reports and used for analysis of sections and core-drillings during fieldwork in east to north greenland from 1982 to 2012 (e.g. the core-drillings blokelv-1, rødryggen-1 and brorson halvø-1; alsen & piasecki 2018; bjerager et al. 2018; alsen et al. 2023). poulsen (1985) analysed a high-resolution series of samples for dinoflagellate cysts across the boundary of the fossilbjerget – hareelv formations at ugleelv. the samples were collected by s. piasecki, and scattered samples from the overlying hareelv formation were sampled by t. birkelund and c. heinberg in ugleelv, jameson land. one ammonite faunal horizon m 9, amoeboceras ilovaiskii subzone, lower a. glosense ammonite zone is correlated with the dinoflagellate cyst stratigraphy (faunal horizon j 49 on jameson land, callomon 1993). as mentioned earlier, the lower black shales assigned to the hareelv formation should have been classified as the olympen formation (see also callomon 1993, fig. 2). upper bathonian to middle oxfordian dinoflagellate cyst assemblages from fossilbjerget and olympen localities on jameson land and from kosmocerasdal on milne land were reported by smelror (1988) based on samples collected by t. birkelund, j.h. callomon and s.  piasecki. the samples were correlated with ammonite stratigraphy, and a dinoflagellate cyst zonation was established up to the middle oxfordian, c. densiplicatum ammonite zone at the last appearance of the dinoflagellate cyst t. scarburghense. gen. et sp. nidarocysta jubilaea was described by monteil (1966) from upper oxfordian – lower kimmeridgian core material in norway and east greenland. the east greenland sample material is from sjællandselv-3 core 303116 in jameson land (bjerager et al. 2018). in a study of stacked sandstone bodies from the bathonian to uppermost oxfordian on milne land and jameson land, larsen et al. (2003) applied integrated ammonite and dinoflagellate cyst stratigraphy, the uppermost part of which overlaps with the present study. kelly et al. (2015) applied jurassic biostratigraphy from north-west europe to east greenland without much documentation. alsen & piasecki (2018) reported integrated ammonite and dinoflagellate cyst from the jurassic blokelv-1 core on jameson land dated to the hareelv formation. relatively few stratigraphical events of dinoflagellate cysts were assigned to the upper oxfordian – lowermost kimmeridgian interval. meanwhile, alsen et al. (2023) reported integrated ammonite and dinoflagellate cyst stratigraphy in the upper jurassic – lower cretaceous cores of the rødryggen-1 and brorson halvø-1 wells drilled in wollaston forland (locations in fig. 1). however, these wells did not reach the oxfordian – lowermost kimmeridgian. 2.2 biozonation palynological analysis of the middle to upper oxfordian sedimentary succession on milne land, east greenland, shows a relatively poor and low diversity dinoflagellate cyst assemblage (fig. 5). however, the assemblage can be referred to the g. jurassica – s. crystallinum zone (gj/sc) of north-west europe (woollam & riding 1983; amended by riding & thomas 1988, 1992). the zone is defined from the last occurrence of t. scarburghense (formerly acanthaulax senta) to the last occurrence of s. crystallinum and is correlated with the base of c. tenuiserratum ammonite zone to the p. baylei ammonite zone, middle oxfordian to lowermost kimmeridgian in northwest europe. https://doi.org/10.34194/geusb.v57.8373 http://www.geusbulletin.org/ piasecki 2024: geus bulletin 57. 8373. https://doi.org/10.34194/geusb.v57.8373 7 of 15 geusbulletin.org in east greenland, the last occurrence of t. scarburghense is also recorded in the lower c. tenuiserratum ammonite zone in kosmoceras dal, milne land, locality 2 (piasecki 1980). the last occurrence of s. crystallinum is recorded in the top of a. rosenkrantzi ammonite zone, locality 39, east of visdal (figs 2–4). the s. crystallinum zone extends the zonation of smelror (1988) to the base of the kimmeridgian. the zone name was simplified to s. crystallinum zone (riding & thomas 1988), and both lower and upper boundaries were subsequently redefined based on other recorded events. the position of the last occurrence of s. crystallinum varies slightly in later reports (riding & thomas 1988, 1992, 1997) from a. rosenkrantzi ammonite zone to p. baylei ammonite zone followed by rare scattered, higher stratigraphic occurrences. the last occurrence of s. crystallinum in the p. baylei ammonite zone is maintained by poulsen & riding (2003). woollam & riding (1983) divided the zone into three subzones. the lower subzone ‘a’ is defined from the last occurrence of t. scharburghense to the last occurrence of compositosphaeridium polonicum and is well recognised in east greenland (fig. 5). the two stratigraphically higher subzones ‘b’ and ‘c’ are not recognised within east greenland stratigraphy. poulsen & riding (2003) applied five subzones, named dsj23–27, to the s. crystallinum zone, including part of the upper t. scarburghense zone (where dsj refers to dinoflagellate cysts, sub-boreal zonation and jurassic). the five subzones are not readily recognised in east greenland partly due to the absence or rarity of the most indicative species. also, the uppermost subzone dsj27 is not recognised in east greenland due to the absence of s. crystallinum in the lowermost kimmeridgian, p. baylei ammonite zone (figs 5 and  6). the dinoflagellate cyst assemblage in the p. baylei ammonite zone is thus referred to the e. luridum zone (woollam & riding 1983; nøhr-hansen 1986) although the nominate dinoflagellate cyst itself is not abundant in east greenland (fig. 6). in contrast, the middle oxfordian to lowermost kimmeridgian dinoflagellate cyst zone interval, named jz32–jz38 in the north atlantic (where jz refers to jurassic zone, sensu bailey 2023), is based on 71 p 23 amoeboceras glosense/serratum el ev at io n 25 20 15 10 5 0 –5 –10 –15 –20 –25 li th os tr at ig ra ph y (* 1 a ld in ge r e lv m b) h al l b re dn in g g ro up (l oc al ity m 3 9) h al l b re dn in g g ro up (l oc al iti es m 2 –5 ) g ro up k ap l es lie f or m at io n fo rm at io n b ay s e lv m em be r *1 k os m oc er es da l m em be r m em be r c hr on os tr at ig ra ph y (* 2 k im m er id gi an ) *2 u pp er o xf or di an m id . o xf od ia n pa rs . su bs ta ge u pp er j ur as si c a m m on ite e ve nt s 33.00 32.00 30.00 24.00 17.00 –7.50 –15.50 –17.25 –25.00 a m m on ite z on at io n 33.00 24.00 17.00 0.00 7.50 17.25 25.50 s .c ry st al lin um z on e (s en su r id in g & t ho m as 1 98 8) s . c ry st al lin um z on e d in of la ge lla te c ys t e ve nt s 32.00 top of s. crystallinum 27.00 base of e. irregulare 13.00 1.00 –2.00 –3.00 –9.50 –14.00 –14.50 base of p. borealis base of h. orbifera top of s. redcliffense base of d. jurassicum top of k. stegasta base of d. minutum top of c. polonicum base of l. subtile base of a. staffinensis base of s. valensii base of s. inritibile top of r. aemulum–16.50 top of c. cerastes –22.50 top of t. scarburghense sa m pl es (g g u s am pl e nu m be rs ) b ar re n stratigraphic range in situ, reworked, '?' occurrences 1 b ar ba ta cy st a pe lio ne ns is 2 b at ia ca sp ha er a la ev ig at a 3 c hy tro ei sp ha er id ia c er as te s 4 c om po si to sp ha er id iu m p ol on ic um 5 d is si lio di ni um h oc ne ra tu m 6 e nd os cr in iu m g al er itu m 7 e nd os cr in iu m lu rid um 8 g on ya ul ac ys ta ju ra ss ic a 9 k al yp te a st eg as ta 10 n an no ce ra to ps is p el lu ci da 11 p ar eo di ni a ce ra to ph or a 12 r hy nc ho di ni op si s cl ad op ho ra 13 r ig au de lla a em ul um 14 r ig au de lla fi la m en to sa 15 s cr in io di ni um c ry st al lin um 16 s irm io di ni um g ro ss ii 17 s ys te m at op ho ra s pp . 18 tr ic ho di ni um s ca rb ur gh en se 19 tu bo tu be re lla a pa te la 20 va le ns ie lla d ic ty di a 21 a to po di ni um h ar om en se 22 s te ph an el yt ro n sc ar bu rg he ns e 23 w re vi tti a? h el ic oi de a 24 e sc ha ris ph ae rid ia p oc oc ki i 25 m en di co di ni um g ro en la nd ic um 26 fr om ea s pp . 27 p ar eo di ni a ha lo sa 28 le pt od in iu m m ira bi le 29 le pt od in iu m s ub til e 30 m ei ou ro go ny au la x pl an os ep ta ta 31 a m bo no sp ha er a st af fin en si s 32 s cr in io di ni um in rit ib ile 33 s te ph an el yt ro n ca yt on en se 34 s ys te m at op ho ra v al en si i 35 b ar ba ta cy st a pi lo sa 36 tu bo tu be re lla e ge m en ii 37 p ar eo di ni a ce ra to ph or a sc op ae a 38 c rib ro pe rid in iu m g ra nu lig er um 39 d in go di ni um m in ut um 40 g on ya ul ac ys ta a de ct a 41 k al yp te a di ce ra s 42 p ro lix os ph ae rid iu m g ra nu lo su m 43 s te ph an el yt ro n re dc lif fe ns e 44 ta en io ph or a iu nc tis pi na 45 b at ia ca sp ha er a pi lo su m 46 d in go di ni um ju ra ss ic um 47 d in go di ni um tu be ro su m 48 e pi pl os ph ae ra c f. sa tu rn al is 49 b at ia ca sp ha er a sp p. 50 e nd os cr in iu m s pp . 51 g lo ss od in iu m d im or ph um 52 li es be rg ia a ff. li es be rg en si s 53 m en di co di ni um s pp . 54 p ar eo di ni a sp p. 55 p ilo si di ni um m yr ia tri ch um 56 s irm io di ni op si s or bi s 57 s ys te m at op ho ra a re ol at a 58 g on ya ul ac ys ta d ua lis 59 h ys tri ch os ph ae rin a or bi fe ra 60 p ar ag on ya ul ac ys ta b or ea lis 61 va le ns ie lla s pp . 62 c hl am yd op ho re lla s pp . 63 c rib ro pe rid in iu m s pp . 64 p ilo si di ni um s p. d f en so m e 19 79 65 k or ys to cy st a pa ch yd er m a 66 s en tu si di ni um s pp . 67 e nd os cr in iu m ir re gu la re 68 e pi pl os ph ae ra g oc ht ii 69 e pi pl os ph ae ra b ire tic ul at a 70 va le ns ie lla o vu lu m p an da di ni um s in os um 31 a m bo no sp ha er a st af fin en si s 21 a to po di ni um h ar om en se 1 b ar ba ta cy st a pe lio ne ns is 35 b ar ba ta cy st a pi lo sa 2 b at ia ca sp ha er a la ev ig at a 45 b at ia ca sp ha er a pi lo su m 49 b at ia ca sp ha er a sp p. 62 c hl am yd op ho re lla s pp . 3 c hy tro ei sp ha er id ia c er as te s 4 c om po si to sp ha er id iu m p ol on ic um 38 c rib ro pe rid in iu m g ra nu lig er um 63 c rib ro pe rid in iu m s pp . 46 d in go di ni um ju ra ss ic um 39 d in go di ni um m in ut um 47 d in go di ni um tu be ro su m 5 d is si lio di ni um h oc ne ra tu m 6 e nd os cr in iu m g al er itu m 67 e nd os cr in iu m ir re gu la re 7 e nd os cr in iu m lu rid um 50 e nd os cr in iu m s pp . 69 e pi pl os ph ae ra b ire tic ul at a 68 e pi pl os ph ae ra g oc ht ii 48 e pi pl os ph ae ra c f. sa tu rn al is 24 e sc ha ris ph ae rid ia p oc oc ki i 26 fr om ea s pp . 51 g lo ss od in iu m d im or ph um 40 g on ya ul ac ys ta a de ct a 58 g on ya ul ac ys ta d ua lis 8 g on ya ul ac ys ta ju ra ss ic a 59 h ys tri ch os ph ae rin a or bi fe ra 41 k al yp te a di ce ra s 9 k al yp te a st eg as ta 65 k or ys to cy st a pa ch yd er m a 28 le pt od in iu m m ira bi le 29 le pt od in iu m s ub til e 52 li es be rg ia a ff. li es be rg en si s 30 m ei ou ro go ny au la x pl an os ep ta ta 25 m en di co di ni um g ro en la nd ic um 53 m en di co di ni um s pp . 10 n an no ce ra to ps is p el lu ci da 71 p an da di ni um s pi no su m 60 p ar ag on ya ul ac ys ta b or ea lis 11 p ar eo di ni a ce ra to ph or a 37 p ar eo di ni a ce ra to ph or a sc op ae a 27 p ar eo di ni a ha lo sa 54 p ar eo di ni a sp p. 55 p ilo si di ni um m yr ia tri ch um 64 p ilo si di ni um s p. d f en so m e 19 79 42 p ro lix os ph ae rid iu m g ra nu lo su m 12 r hy nc ho di ni op si s cl ad op ho ra 13 r ig au de lla a em ul a 14 r ig au de lla fi la m en to sa 15 s cr in io di ni um c ry st al lin um 32 s cr in io di ni um in rit ib ile 66 s en tu si di ni um s pp . 56 s irm io di ni op si s or bi s 16 s irm io di ni um g ro ss ii 33 s te ph an el yt ro n ca yt on en se 43 s te ph an el yt ro n re dc lif fe ns e 22 s te ph an el yt ro n sc ar bu rg he ns e 57 s ys te m at op ho ra a re ol at a 17 s ys te m at op ho ra s pp . 34 s ys te m at op ho ra v al en si i 44 ta en io ph or a iu nc tis pi na 18 tr ic ho di ni um s ca rb ur gh en se 19 tu bo tu be re lla a pa te la 36 tu bo tu be re lla e ge m en ii 20 va le ns ie lla d ic ty di a 70 va le ns ie lla o vu lu m 61 va le ns ie lla s pp . w re vi tti a? h el ic oi de a 234219-tb12 ? 234220-tb11 r ? 234221-10/a 234222-tb/ 10b 234224-tb9 ? 234223-tb10/c 234225-tb8a ? ? 234226-tb8/b ? 234227-tb7 234228-tb6 ? ? 234229-tb4 r ? ? 234230-tb3 234231-tb2 ? 234232-tb1 234177 ? 234176 234175 ? 234171 ? ?234168 234167 ? 234166 ? 234165 234164 234163 234162 234160 ? 234159 r 234155 ? ? ? 234154 ? 234242 234241 ? 234239 234238 dinoflagellate cysts 30 s er ie s pictonia baylei amoeboceras rosenkrantzi amoeboceras rosenkrantzi amoeboceras rosenkrantzi amoeboceras regulare amoeboceras glosense amoeboceras ilovaiskii cardioceras tenuiserratum a. rosenkrantzi zone p. baylei zone a. regulare zone a. serratum – upper a. glosense zones upper a. glosense zone a. glosense zone / a. ilovaiskii subzone c. tenuiserratum zone top of g. adecta fig. 5 dinoflagellate cyst range-chart based on data from a composite succession of kosmocerasdal member to bays elv member, kap leslie formation. the range-chart is combined from successions at localities 2–5 and 39. to fit to the page, the composite succession of kosmocerasdal member is reduced 4 times in thickness and arranged from 0 to –25 m. the succession from locality 39 is shown to scale and spans 34 m from the basis of kosmocerasdal member to the top of bays elv member. abbreviations: mid.: middle. *1: aldinger elv mb. *2: kimmeridgian. dinoflagellate cyst names in figs 5–7. https://doi.org/10.34194/geusb.v57.8373 http://www.geusbulletin.org/ piasecki 2024: geus bulletin 57. 8373. https://doi.org/10.34194/geusb.v57.8373 8 of 15 geusbulletin.org a b c d e f g h i j k m n o p 25 µm q r l fig. 6 stratigraphically significant dinoflagellate cysts from the middle to upper oxfordian s. crystallinum zone in milne land, from locality 39, east of visdal, and localities 2 (kosmoceras dal), 3 (nordøstelv) and 4 (‘hystrix’ dal), eastern milne land (locations in fig. 1). the illustrated specimens are referred to locality numbers, ggu sample number, slide number and england finder coordinates. the illustrated specimens are marked with a red circle on the original slides. magnification is x400 as indicated by a 25 µm scalebar in panel a that applies to all figures. a: scriniodinium crystallinum, locality 39, sample ggu234229, slide 7, e.f. u35/2. b: scriniodinium crystallinum, locality 39, sample ggu234229, slide 7, e.f. p25/2. c: trichodinium scarburghense, locality 2, sample ggu234235, slide 4, e.f. k29/4. d: trichodinium scarburghense, locality 2, sample ggu234235, slide 4, t34/3. e, f: atopodinium haromense, high and low focus, locality 3, figure 6 continued on next page https://doi.org/10.34194/geusb.v57.8373 http://www.geusbulletin.org/ piasecki 2024: geus bulletin 57. 8373. https://doi.org/10.34194/geusb.v57.8373 9 of 15 geusbulletin.org figure 6 continued sample ggu234242, slide 8, p49/2. g: chytroeisphaeridia cerastes, locality 2, sample ggu234155, slide 7, e.f. j26. h: chytroeisphaeridia cerastes, locality 2, sample ggu234155, slide 7, e.f. k52/3. i: rigaudella aemula, locality 2, sample ggu234240, slide 6, e.f. h51. j: ambonosphaera staffinensis, locality 39, sample ggu234227, slide 7, e.f. o26. k: ambonosphaera staffinensis, locality 39, sample ggu234223, slide 7, e.f. u53/2. l: systematophora valensii, locality 39, sample ggu234230, slide 3, e.f. u48/1. m, n: leptodinium subtile, high and low focus, locality 39, sample ggu234219, slide 3, p53/2. o: scriniodinium inritibile, locality 39, sample ggu 234230, slide 7, e.f. r51. p: scriniodinium inritibile, locality 39, sample ggu234220, slide 7, e.f. n28/1. q: compositosphaeridium polonicum, locality 4, sample ggu234159, slide 8, e.f. k36/3. r: compositosphaeridium polonicum, locality 2, sample ggu234239, slide 4, e.f. h46/m4. events with good correlation to events in the s. crystallinum zone in east greenland. some events are based on informally identified species that are not identified in east greenland. the jurassic dinoflagellate cyst stratigraphy of the circum arctic region (bujak et al. 2022) adopted the stratigraphy for lower oxfordian of east greenland of smelror (1988) and reported no data for the middle and upper oxfordian. to preserve a clear and practical definition of the s. crystallinum zone in east greenland, the older definition of the zonal boundaries (woollam & riding 1983) is maintained here. this also provides a simple extension of the preceeding dinoflagellate biostratigraphy of smelror (1988; fig. 6). dinoflagellate cyst events of the s. crystallinum zone were recognised in two other milne land localities in the northern visdal region (larsen et al. 2003, their fig. 1, localities 5, 8 and 12) but were not discussed further in that study. the s. crystallinum zone is also recognisable in the blokelv-1 core on jameson land based on the last occurrence of t. scarburghense to the last occurrence of s. crystallinum (alsen & piasecki 2018, fig. 8). the upper part of the zone is also identified in the eastern slope of tværdal, geographical society ø, in a sandy succession, which is now correlated with the a. regulare – a. rosenkrantzi ammonite zones (surlyk et al. 2023). the s. crystallinum zone is commonly recognised in the oxfordian in both published (piasecki et al. 2004; piasecki & stemmerik 2004) and unpublished materials from the hold with hope and wollaston forland regions but becomes uncertain in more northerly successions from store koldewey to peary land, north greenland (s. piasecki, unpublished data). smelror (2021) analysed palynological material from three boreholes in the ramså basin, andøya, norway. part of the upper bonteigen member, ramså formation, is referred to the p. baylei chronozone, lowermost kimmeridgian, based on dinoflagellate cysts. no oxfordian succession is recorded below this. several dinoflagellate cyst species in the three lowermost sedimentary samples range from the upper oxfordian into the lowermost kimmeridgian in east greenland. however, correlated with data in the present paper, the presence of dingodinium minutum supports smelror’s stratigraphical interpretation of these strata as lowermost kimmeridgian, as this species occurs no higher than the p. baylei ammonite zone in east greenland. 3 results the dinoflagellate cyst assemblage is relatively poor or limited in all samples from the succession at localities 2–5 and 39 (fig. 4). this is partly due to low organic content in the sandy lithology, but, in general, reflects the low abundance and diversity in most of the midto upper oxfordian interval in east greenland. a typical succession commonly contains at least some samples with a few common species. in decreasing abundance, these are g. jurassica, r. cladophora, sirmiodinium grossii, systematophora spp., taeniophora iunctispina and cribroperidinium spp. ten stratigraphical units based on dinoflagellate cyst events within the s. crystallinum zone on milne land are identified in the dinoflagellate cyst range chart (fig. 5). here, the focus is on describing correlative events instead of defining subzones that become less useful north and south of the jameson land basin. specimens mentioned in the following are illustrated in figs 6 and 7. the lowest occurrence of the index species of the s. crystallinum zone (fig. 6a–b) occurs far below the zone itself. it is recorded in the peltoceras athleta ammonite zone, upper callovian, in kosmocerasdal (locality 2, piasecki 1980) as well as in the p. athleta ammonite zone in north-west europe (riding & thomas 1992). the highest occurrence of t. scarburghense (figs 5 and 6c–d) defines the lower boundary of the s. crystallinum zone in the lower c. tenuiserratum ammonite zone, in kosmocerasdal (locality 2, piasecki 1980). the lowest occurrences of glossodinium dimorphum and scriniodinium inritibile were implemented by poulsen & riding (2003) into the definition of the basal boundary of the s. crystallinum zone as lower than the highest occurrence of t. scarburghense in north-west europe. however, g. dimorphum is very rare, and like s. inritibile, it appears in the a. glosense ammonite zone in east greenland. atopodinium haromense (fig. 6e–f) with vague morphological characteristics occurs rarely in the lowermost s. crystallinum https://doi.org/10.34194/geusb.v57.8373 http://www.geusbulletin.org/ piasecki 2024: geus bulletin 57. 8373. https://doi.org/10.34194/geusb.v57.8373 10 of 15 geusbulletin.org b c d e f g h i j k m n o 25 µm a 25 µm p s rq l t u v 25 µm fig. 7 stratigraphically significant dinoflagellate cysts from the middle to upper oxfordian in milne land from locality 39, kosmoceras dal, east of visdal, and localities 2 and 6, cardioceraskløft, in eastern milne land (locations in fig. 1). the illustrated specimens are referred to locality numbers, ggu figure 7 continued on next page https://doi.org/10.34194/geusb.v57.8373 http://www.geusbulletin.org/ piasecki 2024: geus bulletin 57. 8373. https://doi.org/10.34194/geusb.v57.8373 11 of 15 geusbulletin.org zone but becomes more morphologically characteristic upwards. smelror (1988) refers the oxfordian atopodinium species with no atopodinium prostatum characters to morphological variation of an a. prostatum-complex, whereas riding & thomas (1997) record a. haromense from the basal middle oxfordian and upwards. chytroeisphaeridia cerastes (fig. 6g–h) occurs with certainty up to the lower a. glosense ammonite zone and the a. ilovaiskii subzone. morphologically similar specimens occur higher in the succession but both their small size and uncertain archaeopyle type suggest another classification for those specimens. rigaudella aemula (fig. 6i) commonly occurs up to the lower a. glosense ammonite zone and a. ilovaiskii subzone. leptodinium subtile (fig. 6m–n) and leptodinium mirabile appear contemporaneously in this horizon, but only l. subtile occurs regularly in higher strata. ambonosphaera staffinensis (fig. 6j–k), s. inritibile (fig. 6o–p) and systematophora valensii (fig. 6l) appear from this horizon and continue above the s. crystallinum zone. compositosphaeridium polonicum (fig. 6q–r) occurs consistently but not commonly, with a highest occurrence in the lower a. glosense ammonite zone and a. ilovaiskii subzone. the lowermost occurrence of d. minutum (fig. 7a–e) appears in a. glosense ammonite zone and a. ilovaiskii subzone. specimens of d. minutum may represent morphological variants of a. staffinensis (fig. 6j–k) since specimens occur with apparent transitional morphology. however, d. minutum has a prominent tabulation, an antapical keel with a vertical antapical plate on the ventral side and a prominent flagellar scar in the mid-sulcal area. d. minutum was previously informally recorded as a. ‘utera’ by piasecki (1980). kalyptea stegasta (fig. 7f–g) is common and has a clear uppermost occurrence in the upper a. glosense ammonite zone. it is recorded higher neither in the s. crystallinum zone nor above, but it is recorded to the top of the lower volgian in north-west europe (riding & thomas 1992). dingodinium jurassicum (fig. 7h) is present in the uppermost sample from the a. glosense ammonite zone of kosmocerasdal member and the three lowermost samples of locality 39 in the upper a. glosense/a. serratum ammonite zone. the short range of this species apparently forms a stratigraphically narrow and locally useful occurrence. a few instances of dingodinium tuberosum (fig. 7i) are also recorded in the lowermost level with d. jurassicum. this deviates from north-west europe, where the first occurrences of d. jurassicum and d. tuberosum are recorded at stratigraphically different levels: lower and higher, respectively (e.g. riding & thomas 1992; poulsen 1996). the sculpture of the endophragm and the overall shape and morphology of d. jurassicum are quite different from d. tuberosum (fig. 7h–i). hystrichosphaerina orbifera (fig. 7l) appears in the upper a. glosense/a. serratum ammonite zones. chorate cysts are rare in most samples and mostly too crushed or degraded for reliable identification. in comparison, the characteristic process complexes of h. orbifera make it an identifiable species. paragonyaulacysta borealis (fig. 7o–p) also appears in this sample. in east greenland, p. borealis is obviously tabulated but with vague parasutures. hence, specimens are mostly identified on the basis of overall size, shape, surface sculpture of the autophragm and the apicular structure in comparison with more distinct tabulated specimens in north-east and north greenland. gonyaulacysta adecta (fig. 7q–r) occurs highest in the a. rosenkrantzi ammonite zone. g. adecta is basically a g. jurassica but lacking a hypocoel. all morphological structure on g. jurassica varies radically possibly in response to environmental or climatic conditions. endoscrinium irregulare (fig. 7s–v) appears in the same sample level. e. irregulare is not considered stratigraphically significant, but it has a significant morphology and a clear first occurrence in the a. rosenkrantzi ammonite zone. scriniodinium crystallinum (fig. 6a–b) has its highest occurrence in the top of the a. rosenkrantzi ammonite figure 7 continued sample numbers, slide number and england finder coordinates. the illustrated specimens are marked with a red circle on the original slide. a: dingodinium minutum, locality 6, sample ggu245827, slide 8, e.f. g32/2. magnified x600. b: dingodinium minutum, locality 6, sample ggu245828, slide 8, e.f. s53/3. magnified x600. scale in a. c, d: dingodinium minutum, high and low focus, locality 6, sample ggu245830, slide 9, e.f. j47. magnified x600. scale in a. e: dingodinium minutum, locality 6, sample ggu245830, slide 9, e.f. h29. magnified x600. scale in a. f: kalyptea stegasta, locality 2, sample ggu234159, slide 8, e.f. t27/4. magnified x400. g: kalyptea stegasta, locality 2, sample ggu234239, slide 4, e.f. g34. magnified x400. scale in f. h: dingodinium jurassicum, locality 39, sample ggu234231, slide 9, u37/2. magnified x400. scale in f. i: dingodinium tuberosum, locality 39, sample ggu234231, slide 6, g23/3. magnified x400. scale in f. j, k: stephanelytron redcliffense, high and low focus, locality 39, sample ggu234232, slide 3, e.f. m27/1. magnified x400. scale in f. l: systematophora areolata, locality 39, sample ggu234230, slide 3, e.f. v37/1. magnified x400. scale in f. m: hystrichosphaerina orbifera, locality 39, sample ggu234221, slide 3, e.f. o49/2. magnified x400. scale in f. n: hystrichosphaerina orbifera, locality 39, sample ggu234220, slide 3, e.f. n45/2. magnified x400. scale in f. o: paragonyaulacysta borealis, locality 39, sample ggu234223, slide 6, e.f. o39/3. magnified x400. scale in f. p: paragonyaulacysta borealis, locality 6, sample ggu245827, slide 8, e.f. r32/3-s32/1. magnified x400. scale in f. q: gonyaulacysta adecta, locality 39, sample ggu234228, slide 10, e.f. j48/2. magnified x400. scale in f. r: gonyaulacysta adecta, locality 39, sample ggu234232, slide 3, e.f. p23. magnified x400. scale in f. s: endoscrinium irregulare, locality 39, sample ggu234224, slide 7, e.f. t28/3–4. magnified x400. scale in f. t: endoscrinium irregulare, locality 39, sample ggu234220, slide 7, e.f. s48. magnified x400. scale in f. u: endoscrinium irregulare, locality 39, sample ggu234220, slide 6, e.f. n45.4. magnified x400. scale in f. v: endoscrinium irregulare, locality 6, sample ggu245827, slide 8, e.f. w35. magnified x600. https://doi.org/10.34194/geusb.v57.8373 http://www.geusbulletin.org/ piasecki 2024: geus bulletin 57. 8373. https://doi.org/10.34194/geusb.v57.8373 12 of 15 geusbulletin.org zone at locality 39. rare specimens have been recorded much higher in the succession at other localities, but they are considered to have been reworked or are possibly misinterpretations. dingodinium minutum (fig. 7a–e) occurs highest into the p. baylei ammonite zone, lowermost kimmeridgian. this is higher than the s. crystallinum zone and the lowermost e. luridum zone. ammonite zones p. baylei d. minutum a. rosenkrantzii w an ae a fim br ia ta a. regulare a. serratum a. glosense c. tenuiserratum c. densiplicatum c. cordatum q. mariae (q. lamberti?) p. athleta range of index species events lads fads o xf or di an k im m er id ia n d in of la ge lla te cy st z on es s . c ry st al lin um t. s ca rb ur gh en se w . f im br ia ta c al lo vi an tr ic ho di ni um s ca rb ur gh en se s cr in io di ni um c ry st al lin um e nd os cr in iu m lu rid um g. adecta s. redcliffense k. stegasta c. polonicum / r. aemula t. scarburghense w. fimbriata lad: last apperance datum s. crystallinum e. irregulare p. borealis / h. orbifera d. jurassicum s. crystallinum w. fimbriata t. scarburghense fad: first apperance datum legend d. minutum * * e. luridum l. subtile a. staffinensis s. valensii s. inritibile lo w er m id dl e u pp er fig. 8 schematic correlation of the oxfordian ammonite zones, dinoflagellate cyst zones and events on milne land, east greenland. https://doi.org/10.34194/geusb.v57.8373 http://www.geusbulletin.org/ piasecki 2024: geus bulletin 57. 8373. https://doi.org/10.34194/geusb.v57.8373 13 of 15 geusbulletin.org 4 discussion the s. crystallinum zone appears to be a useful identifier of middle to upper oxfordian strata in east greenland based on palynological analyses (fig. 8). seventeen dinoflagellate cyst events characterise the lower and upper boundaries and subdivide the zone into 10 units. the dinoflagellate cyst zone correlates with five succeeding ammonite zones, from lowermost to uppermost: the c. tenuiserratum, a. glosense, a. serratum, a. regulare and a. rosenkrantzi zones. the stratigraphical positions of both base and top of the s. crystallinum zone in the study area deviate slightly from the corresponding zone in north-west europe, but this may reflect the limited amount or spacing of samples studied in east greenland. barski (2018) shows that common s. crystallinum reaches into the lower p. baylei ammonite zone on the isle of skye, scotland. however, other dinoflagellate cyst studies in north-west europe show different stratigraphical ranges to east greenland, such as perisseiasphaeridium pannosum and senoniasphaera jurassica. most of the events in north-west europe applied for stratigraphical subdivision of zones into subzones do not appear at the same stratigraphical levels in east greenland. the lowermost appearance of t. scarburghense is between two faunal horizons of m 2 in kosmocerasdal and is referred to the p. athleta ammonite zone (fig. 4). the lower boundary of the w. fimbriata zone (smelror 1988) was defined by the first occurrence of w. fimbriata below faunal horizon m 3, q. mariae ammonite zone. the assumed q. lamberti ammonite zone beneath faunal horizon m 3 is not recorded in the kosmocerasdal succession (locality 2), but the lower boundary of the w. fimbriata zone is still hypothetically referred to this ammonite zone marked by (?) in fig. 8. the lower boundary of the t. scarburghense zone (smelror 1988) was defined by the highest occurrence of w. fimbriata, now referred to the lower cardioceras cordatum ammonite zone (fig. 8). the lower boundary of the s. crystallinum zone in east greenland coincides with the highest occurrence of t. scarburghense zone (smelror 1988) defined by the highest occurrence of t. scarburghense in the c. tenuiserratum ammonite zone (fig. 8). the upper zonal boundary is in the uppermost a. rosenkrantzi ammonite zone and not the p. baylei ammonite zone as in north-west europe. some recorded species may have a slightly problematic taxonomic affinity due to vaguely expressed morphological characters or bad preservation. parago nyaulacysta borealis has very discrete tabulation on milne land, but it is recorded consistently from jameson land towards the north along east and north-east greenland to north greenland (håkansson et al. 1981). a. staffinensis and d. minutum appear to have transitional morphologies, and the distinction between the two may be uncertain in some cases. some c. cerastes-shaped specimens have been excluded here due to their small size and incomplete archaeopyles, and many systematophora-like specimens cannot be identified to species level due to fragmentation, folding or bad preservation. epiplosphaera saturnalis is not typically recorded here, but the reduced surface sculptural elements on e. cf. saturnalis are similar to the elements of the holotype (brideaux & fisher 1976; plate 6, figs 1–7 and plate 7, fig.  10). comparable specimens have been recorded from milne land to north greenland as lanterna saturnalis (håkansson et al. 1981). 5 conclusions in this study, the history of oxfordian dinoflagellate cyst stratigraphy in east greenland is summarised and updated. the dinoflagellate cysts stratigraphy of the marine deposited, middle to upper jurassic succession is documented and correlated with ammonite stratigraphy. it is found that the s. crystallinum zone from northwest europe is applicable in east greenland in its original definition. however, the later associated stratigraphical markers for both the lower and upper boundary are not directly applicable in east greenland. the subdivision of five subzones in north-west europe is mostly based on species that are rare or absent in the east greenland, including the first and last subzones (djs23 and -27), which are not included in the s. crystallinum zone in east greenland. fifteen events are suggested to subdivide the s. crystallinum zone into 10 units by characteristic and common species in east greenland. the zonal boundaries and the stratigraphic events are directly integrated with the ammonite stratigraphy. the s. crystallinum zone is applicable in the jurassic sedimentary basins of east and north-east greenland and may be applicable in parts of the north atlantic region. with this contribution to the jurassic biostratigraphy in east greenland and together with the upper jurassic biostratigraphy of two previously drilled cores rødryggen-1 and brorson halvø-1, the whole jurassic succession in east greenland has now been palynologically analysed based on spores and pollen or dinoflagellate cysts. acknowledgements kim villadsen, palynological laboratory, prepared the analysed samples. the present manuscript was compiled as emeritus at the globe institute, university of copenhagen. staff in the department of geophysics and sedimentary basins, geological survey of denmark and greenland (geus), helped draw the figures. peter alsen is thanked for helpful criticism of the manuscript and the results. the reviewers j.b. riding and morten smelror are thanked for helpful corrections of the text and suggested improvements of the content. https://doi.org/10.34194/geusb.v57.8373 http://www.geusbulletin.org/ piasecki 2024: geus bulletin 57. 8373. https://doi.org/10.34194/geusb.v57.8373 14 of 15 geusbulletin.org additional information funding statement the original fieldwork and office studies were financed by danish natural science research foundation grants in 1977–1980. conflict of interests the author declares no competing interests. author contributions sp: conceptualisation; investigation; visualisation; writing – original draft; writing – review and editing. references alsen, p. & piasecki, s. 2018: biostratigraphy of the hareelv formation (upper jurassic) in the blokelv-1 core, jameson land, central east greenland. geological survey of denmark and greenland bulletin 42, 15–37. https://doi.org/10.34194/geusb.v42.4308 alsen, p., piasecki, s., nøhr-hansen, h., pauly, s., sheldon, e. & hovikoski, j. 2023: stratigraphy of the upper jurassic to lowermost cretaceous in the rødryggen-1 and brorson halvø-1 boreholes, wollaston forland, ne greenland. geus bulletin 55, 8342. https://doi.org/10.34194/geusb.v55.8342 bailey, d. 2023: biostratigraphic consultancy company. biostrat ltd. http://www.biostrat.org.uk/ (accessed 10 june 2023) barski, m. 2018: dinoflagellate cyst assemblages across the oxfordian/ kimmeridgian boundary (upper jurassica) at flodigarry, staffin bay, isle of skye, scotland – a proposed gssp for the base of the kimmeridgian. volumina jurassica 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sarjeant, w.a.s. 1972: dinoflagellate cysts and acritarchs from the upper vardekløft formation (jurassic) of jameson land, east greenland. meddelelser om grønland 195(4), 64 pp. smelror, m. 1988: late bathonian to early oxfordian dinoflagellate cysts stratigraphy of jameson land and milne land, east greenland. grønlands geologiske undersøgelse rapport 137, 135–159. https://doi. org/10.34194/rapggu.v137.8019 smelror, m. 2021: palynostratigraphy, palynofacies, t-r cycles and paleoenvironments in the middle jurassic–early cretaceous ramså basin, andøya, northern norway. geosciences 11, 354. https://doi. org/10.3390/geosciences11090354 surlyk, f. et al. 2021: jurassic stratigraphy of east greenland. geological survey of denmark and greenland bulletin 46, 6521. https://doi. org/10.34194/geusb.v46.6521 surlyk, f., alsen, p., hovikoski, j. & piasecki, s. 2023: uplift, deflation and marine onlap of a jurassic rift dome, illustrated by a backstepping middle–upper jurassic shelf-to-slope succession, geographical society ø, east greenland. terra nova 35, 506–513. https://doi. org/10.1111/ter.12673 surlyk, f., callomon, j.h., bromly, r.g. & birkelund, t. 1973: stratigraphy of the jurassic–lower cretaceous sediments of jameson land and scoresby land, east greenland. grønlands geologiske undersøgelse, bulletin 105, 94 pp. https://doi.org/10.34194/bullggu.v105.6646 sykes, r.m. & surlyk, f. 1976: a revised ammonite zonation of the boreal oxfordian and its application in northeast greenland. lethaia 9, 421–436. https://doi.org/10.1111/j.1502-3931.1976.tb00984.x woollam, r. & riding, j.b. 1983: dinoflagellate cyst zonation of the english jurassic. institute of geological sciences report 83(2), 1–42. https://doi.org/10.34194/geusb.v57.8373 http://www.geusbulletin.org/ https://doi.org/10.34194/rapggu.v137.8019 https://doi.org/10.34194/rapggu.v137.8019 https://doi.org/10.3390/geosciences11090354 https://doi.org/10.3390/geosciences11090354 https://doi.org/10.34194/geusb.v46.6521 https://doi.org/10.34194/geusb.v46.6521 https://doi.org/10.1111/ter.12673 https://doi.org/10.1111/ter.12673 https://doi.org/10.34194/bullggu.v105.6646 https://doi.org/10.1111/j.1502-3931.1976.tb00984.x the scriniodinium crystallinum dinoflagellate cyst zone in the middle-upper oxfordian, upper jurassi 1 introduction 2 material and methods 2.1 history of dinoflagellate cyst stratigraphy in east greenland 2.2 biozonation 3 results 4 discussion 5 conclusions acknowledgements additional information funding statement conflict of interests author contributions references figures fig. 1 map of the study area. a: geological map of east greenland (modified from surlyk et al. 2021) with locations of the rødryggen-1 and brorson halvø-1 drilling sites. inset: maps of localities analysed in this study. b: jameson land with blokelv-1 core and c: ilimananngip nunaa (milne land) with localities m 2, 3, 4, 5, 6 and 39 (locality numbers sensu birkelund et al. 1984). fig. 2 lithostratigraphical scheme of jurassic to lower cretaceous sediments on milne land based on callomon & birkelund (1980), birkelund & callomon (1985), birkelund et al. (1984) and surlyk et al. (2021). fig. 3 ammonite faunal horizons on milne land correlated with ammonite zones (callomon & birkelund 1980; birkelund & callomon 1985). shading indicates the faunal province affiliations. fig. 4 measured and sampled sedimentary successions from milne land with ammonite faunal horizons (e.g. m 14) and palynological samples (e.g. ggu245951) modified from piasecki (1980). the locality numbers are from piasecki (1980, fig.3) and birkelund & callomon (1985, fig. 2). the sedimentary succession from localities 2–5 shows the complete kosmocerasdal member, kap leslie formation. sedimentary logs from locality 6, cardioceraskløft, and locality 39, east of visdal, are correlated by the upper boundary of aldinger elv member and the basis of bays elv member. fig. 5 dinoflagellate cyst range-chart based on data from a composite succession of kosmocerasdal member to bays elv member, kap leslie formation. the range-chart is combined from successions at localities 2–5 and 39. to fit to the page, the composite succession of kosmocerasdal member is reduced 4 times in thickness and arranged from 0 to –25 m. the succession from locality 39 is shown to scale and spans 34 m from the basis of kosmocerasdal member to the top of bays elv member. abbreviations: mid.: middle. *1: aldinger elv mb. *2: kimmeridgian. dinoflagellate cyst names in figs 5–7. fig. 6 stratigraphically significant dinoflagellate cysts from the middle to upper oxfordian s. crystallinum zone in milne land, from locality 39, east of visdal, and localities 2 (kosmoceras dal), 3 (nordøstelv) and 4 (‘hystrix’ dal), eastern milne land (locations in fig. 1). the illustrated specimens are referred to locality numbers, ggu sample number, slide number and england finder coordinates. the illustrated specimens are marked with a red circle on the original slides. magnification is x400 as indicated by a 25 μm scalebar in panel a that applies to all figures. a: scriniodinium crystallinum, locality 39, sample ggu234229, slide 7, e.f. u35/2. b: scriniodinium crystallinum, locality 39, sample ggu234229, slide 7, e.f. p25/2. c: trichodinium scarburghense, locality 2, sample ggu234235, slide 4, e.f. k29/4. d: trichodinium scarburghense, locality 2, sample ggu234235, slide 4, t34/3. e, f: atopodinium haromense, high and low focus, locality 3, fig. 7 stratigraphically significant dinoflagellate cysts from the middle to upper oxfordian in milne land from locality 39, kosmoceras dal, east of visdal, and localities 2 and 6, cardioceraskløft, in eastern milne land (locations in fig. 1). the illustrated specimens are referred to locality numbers, ggu sample numbers, slide number and england finder coordinates. the illustrated specimens are marked with a red circle on the original slide. a: dingodinium fig. 8 schematic correlation of the oxfordian ammonite zones, dinoflagellate cyst zones and events on milne land, east greenland. research article | short voutchkova 2023: geus bulletin 53. 8352. https://doi.org/10.34194/geusb.v53.8352 1 of 8 temporal variation of iodine in danish groundwater denitza d. voutchkova* geological survey of denmark and greenland (geus), aarhus, denmark *correspondence: dv@geus.dk received: 15 may 2023 revised: 18 july 2023 accepted: 20 july 2023 published: 18 aug 2023 keywords: iodine, temporal variation, groundwater, denmark abbreviations b.g.l.: below ground level gam: generalised additive model geus: geological survey of denmark and greenland grumo: danish national groundwater monitoring programme icp-ms: inductively coupled plasma mass spectrometry idd: iodine deficiency disorders iqr: interquartile range lod: limit of detection loess: local polynomial regression fit mad: median absolute deviation rcv: robust coefficient of variation u: analytical uncertainty geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: adam hambly (technical university of denmark) reviewed by: andre banning (university of greifswald, germany) and one anonymous reviewer. funding: see page 7 competing interests: see page 7 additional files: see page 7 abstract iodine is an essential element for human health, and both high and low iodine intake could have negative health outcomes. the spatial variation of iodine in danish groundwater has been studied before, but to the author’s knowledge, this is the first time that the temporal variation is characterised. nationwide data from the danish groundwater monitoring programme (grumo) were analysed between 2011 and 2021, including 2924 samples from 1242 well screens at 893 wells. the sampling frequency varied and so the robust coefficient of variation (rcv) was calculated for 930 (75%) of well screens, and time-series analysis was performed for 23 (2%). key findings are (1) iodine in danish groundwater varies over time (0–124%, median = 10%), (2) in one quarter of the well screens rcv exceeds 20% and (3) this variation cannot be attributed solely to analytical uncertainty at 14% of the well screens. the impact of temporal variation of iodine in danish drinking water of groundwater origin should be evaluated in future exposure or epidemiological studies with respect to the study goal, location and time period. since the temporal variation could not be quantified over the entire concentration range, monitoring of iodine in danish groundwater should continue. introduction iodine is an essential element for the proper functioning of the thyroid, and either high or low iodine intake could result in adverse human health outcomes, given that there is a u-shaped relationship between iodine intake and thyroid disorders (laurberg et al. 2009) or autoimmunity (wang et al. 2019). iodine deficiency disorders (idd) are a major public health problem globally (de benoist et al. 2004), and even though levels of global iodine nutrition have improved since the 20th century, andersson et al. (2012) estimated that 1.88 billion people, including 241 million school children, still had insufficient dietary iodine intakes. iodine deficiency has been described as “the greatest cause of preventable brain damage in childhood” (de benoist et al. 2004). iodine deficiency could also cause miscarriages or stillbirths (at the foetal stage); neonatal goitre or hypothyroidism, endemic mental retardation (at the neonatal stage); goitre, hypoor hyperthyroidism, impaired mental function or retarded physical development (in children, adolescents and adults; de benoist et al. 2004). excessive iodine intakes, on the other hand, may result in hypoor hyperthyroidism, goitre and/or thyroid autoimmunity for some individuals (farebrother et al. 2019). the recommended daily intake and the tolerable upper iodine intake vary with age (table 1; institute of medicine 2001; who & faoun 2004; efsa 2006; who 2007). drinking water could be a significant contributor to the daily iodine intake for some populations (voutchkova et al. 2014; farebrother et al. 2019; ma et al. 2022). recently, a meta-analysis by azevedo et al. (2023) showed that iodine status is directly correlated to iodine content in drinking water and concluded that iodine concentration in drinking water can be used as an indicator of dietary intake. therefore, the spatial variation of iodine in both drinking https://doi.org/10.34194/geusb.v53.8352 https://orcid.org/0000-0003-2840-072x mailto:dv@geus.dk https://creativecommons.org/licenses/by/4.0/deed.ast voutchkova 2023: geus bulletin 53. 8352. https://doi.org/10.34194/geusb.v53.8352 2 of 8 www.geusbul let in.org water and groundwater and its importance for human health have been discussed in depth in the scientific literature: either with respect to the need for optimising dietary iodine intake of a given population subjected to a universal salt iodisation programme (voutchkova et al. 2014; ma et al. 2022) or because of the potential for the formation of unwanted iodinated by-products when implementing certain advanced methods of drinking water treatment (sharma et al. 2019). however, less attention has been given to temporal aspects of iodine variation, which may be important regarding iodine exposure from drinking water. in denmark, drinking water is of entirely groundwater origin and undergoes mostly simple treatment (aeration and sand filtration) without chlorination or ozonation. it has been hypothesised that iodine concentrations in danish groundwater are stable over time. this hypothesis stems from limited evidence for iodine in treated drinking water (tap water) and was mostly based on qualitative assessments, which concluded the following: 1. day-to-day variation within a period of 10 days at two locations in denmark was “small” (pedersen et al. 1999) 2. “no significant” difference between samples collected in january and june (rasmussen et al. 2000) 3. concentration at a waterworks in the northernmost part of denmark (skagen) was “unaltered” in the period 1997–2000, based on one sample taken every 2 months for a duration of 6 months and one sample taken every year for a period of 4 years (andersen et al. 2002) 4. no conclusive results concerning short-term variation were observed in eight samples taken in a 2-week period at a waterworks supplying copenhagen (voutchkova et al. 2014). the purpose of this article is therefore to quantify the temporal variation of total dissolved iodine in danish groundwater. this assessment is based on nationwide groundwater-quality monitoring data covering a 10-year period, which has not been reported and systematically analysed previously. data total dissolved iodine (‘iodine’ for brevity) was analysed in 2924 samples obtained during the period 2011–2021 (31 january 2011 – 11 may 2021) as part of denmark’s national groundwater monitoring programme grumo (thorling et al. 2023). these samples were taken at 1242 sampling points (well screens) belonging to 893 grumo wells located throughout the country (fig. 1). the well screens were relatively short: 1 or 2 m long (q25 = 1 m, median = 1 m, q75 = 2 m, min. to max.: 0.05–69 m). following the sampling protocols of the geological survey of denmark and greenland (geus), well purging was carried out until ph, conductivity, o2 and temperature were stable (thorling 2012). samples were filtered in the field through 0.45-µm filters and placed in glass bottles without conservation or other sample pre-handling, after which the samples were stored in cold (0–4°c) and dark conditions for a maximum of 36  h before analysis (ministry of environment of denmark 2011). iodine was analysed with icp-ms (ministry of environment of denmark 2011) at nationally accredited laboratories. the national requirements for analytical uncertainty of all environmental measurements for iodine are 1.5 µg/l absolute expanded uncertainty (uabs) for low concentrations and 20% relative expanded uncertainty (urel) for high concentrations (ministry of environment of denmark 2021). as ministry of environment of denmark (2021) has not defined what is considered a low concentration, the highest of the two (max (urel, uabs)) is used here as the analytical uncertainty for individual samples (u). the actual values of u are unknown for this monitoring data set. for 71% of the samples (n = 2082), uabs ≥ urel, so u = uabs. for the remaining 29% (n = 842), urel was higher, so u = urel. fewer than 1% of the samples (n = 26) were below the lod. two different lods were used: 2 µg/l (10 samples) and 0.3 µg/l (16 samples). the values below lod were handled by substitution with 0.5 × lod, which is equal to 1 µg/l or 0.15 µg/l, respectively. statistical methods the sampling frequency varied at the different well screens from 1 to 11 times during the study period (figs 1 and 2a). the temporal variability could only be assessed for sampling locations with at least two samples (n = 930, 75%) using the robust coefficient of variation (rcv, equation 1; arachchige et al. 2022). table 1 recommended daily intake (rdi) and tolerable upper intake (tui) for iodine. age rdia (µg/day) tui (usa)b (µg/day) tui (eu)c (µg/day) children 1201 300–9002 200–5003 adults 150 1100 600 pregnant or lactating women 250 1100 600 for children, the provided values and ranges are for: 1school children; 2children in the age 4–18 years; 3children in the age 1–17 years. references are as follows: aworld health organization 2007; world health organization & food and agriculture organization of the united nations 2004. binstitute of medicine, academy of sciences & usa 2001. cefsa 2006. https://doi.org/10.34194/geusb.v53.8352 http://www.geusbulletin.org voutchkova 2023: geus bulletin 53. 8352. https://doi.org/10.34194/geusb.v53.8352 3 of 8 www.geusbul let in.org the rcv is a measure of relative dispersion similar to the coefficient of variation but is based on the median  and  median absolute deviation (mad) as follows: rcv = 1.4826 × mad/m (1) where m is the median and mad is a robust measure of variability, similar to the standard deviation but using the median instead. mad (equation 2) is defined as the median of the absolute deviations from the median of the data (m), such that: mad = median(|xi – m|) (2) where xi are the individual iodine measurements at a given sampling location. these metrics for central tendency and variability are more robust against outliers or skewed distributions and so were preferred for this study. next, rcv was compared against the analytical uncertainty across denmark to provide an indication as to whether the observed iodine variability might be associated with environmental factors or else most probably because of the analytical uncertainty. this was done as follows: > ×       → ≤ ×       →         rcv u x rcv u x if median 100% environmental median 100% analytical j i i j i i (3) where i is the index of an individual sample, j is the index of the well screen, x is iodine concentration and u is analytical uncertainty. here the analytical uncertainty is converted to a percentage, so that it could be compared to rcvj. this comparison is only used as a screening tool to provide a preliminary estimate of the potential cause of the observed variation. a more comprehensive assessment at a local scale involving  an increased number of samples, as well as additional hydrogeochemical, hydrogeological and environmental data would be needed to provide further details. 0 50 100 km number of iodine samples in grumo well screens (n) 1 2–9 10–11 (labelled well sites) germany 65 sweden 94 96 period: 2011–2021 data extraction: 22 june 2021 105 121 131 141 159 216 n fig. 1 spatial distribution of iodine sampling sites in the study period at well screens of the danish national groundwater monitoring programme (grumo). https://doi.org/10.34194/geusb.v53.8352 http://www.geusbulletin.org voutchkova 2023: geus bulletin 53. 8352. https://doi.org/10.34194/geusb.v53.8352 4 of 8 www.geusbul let in.org in addition, a trend analysis was performed solely for well screens with ≥10 years of data (n = 23; 2%). linear regression and local polynomial regression (loess with a generalised additive model, gam, with  integrated smoothness estimation as implemented in wood (2023)) and their 95% confidence intervals  were used to interpret the trends (wickham 2016). all statistical analyses and summaries were implemented  in r v. 4.2.1 (r core team 2022). results and discussion nationwide temporal variability the results of this nationwide assessment showed that the robust coefficient of variation (rcv) for iodine at  the 930 sampling points in denmark had a wide range (0–124%) with a median of 10% (iqr 4–21%; fig. 2b). this means that at half (quarter) of the sampling locations, iodine variation in the period exceeded 10% (20%). part of the observed variation could be due to fig. 2 iodine in groundwater wells from the danish national groundwater monitoring programme (grumo). a: histogram of number of iodine samples per well screen. b: scatter plot of robust coefficient of variation (rcv) against median iodine concentration, where each symbol refers to a well screen. c: iodine concentration and its variation (expressed by the median and median absolute deviation, mad). grouped according to the number of samples per sampling location, where horizontal jitter is added to minimise overlapping points. c: grey colour is used when the mad was not calculated (only one sample available) and if mad = 0 (2–11 samples). 312 678 110 33 36 16 13 5 12 18 9 iodine samples (n) sa m pl in g po in ts (n ) 0 3 6 9 12 0 200 400 600 study period: 2011–2021 (a) 2 4 6 8 10 samples (n) median iodine concentration (μg/l) 0 1 10 100 500 r ob us t c v (% ) 40 80 120 (b) 1 2 3 4 5 6 7 8 9 10 11 io di ne (m ed ia n ± m a d , μ g/ l) (c) iodine samples per sampling point (n) 0.1 10 100 1000 1 mad (μg/l) 10 1 0.1 0.01 https://doi.org/10.34194/geusb.v53.8352 http://www.geusbulletin.org voutchkova 2023: geus bulletin 53. 8352. https://doi.org/10.34194/geusb.v53.8352 5 of 8 www.geusbul let in.org analytical uncertainty, and so comparison between rcv and u was used as a screening tool at the national scale. this comparison showed that at 14% of the sampling points (n = 127, fig. 3), iodine variability could not be attributed solely to analytical uncertainty. it could be inferred that at those locations, the variability was, at least partially, caused by environmental factors such as local hydrogeological conditions and/ or iodine variability in precipitation. no concentration dependency or depth dependency (not shown here) or spatial pattern (fig. 3) was associated with the observed iodine variability at the national scale. further, more focused investigations could elucidate the governing environmental factors at specific locations. this assessment showed that there is a temporal variability of iodine in groundwater, but its importance in exposure and epidemiologic studies cannot be inferred purely from these data alone. as all drinking water in denmark originates from groundwater, variability in groundwater iodine may have a significant effect on the concentration of iodine in finished (i.e. treated) drinking water, especially in parts of the country where high levels of iodine are observed. other factors could also contribute to iodine variability in drinking water. for example, the well sites, wells and/or the pumping strategies could change in time, resulting in the use of other aquifers or parts of the aquifer where the iodine concentration is different. it is also possible that the water treatment and its performance over time affect the iodine content of the finished product. therefore, the recommendation is to consider temporal aspects when designing future studies, so that the significance of iodine variability can be assessed with respect to the specific location of interest. even though the grumo programme has yielded plenty of new iodine data over the past decade, it is still challenging to assess variability for the full range of observed concentrations. the highest iodine concentrations (>50 µg/l) and absolute variations occur at locations with only two or very few samples (fig. 2c), which limits analysis solely to a rcv to u comparison. at 25% of the sampling locations (n = 312), only a single sample was acquired, and the variability could therefore not be assessed at those sites. it is therefore recommended that iodine continues to be monitored as part of the grumo programme. it should be noted that iodine is not analysed in drinking water in denmark because fig. 3 iodine variability attribution at sampling sites throughout denmark. n 0 50 100 km iodine variability attribution analytical uncertainty (n = 803) environmental factors (n = 127) not enough data (n = 312) germany sweden period: 2011–2021 data extraction: 22 june 2021 https://doi.org/10.34194/geusb.v53.8352 http://www.geusbulletin.org voutchkova 2023: geus bulletin 53. 8352. https://doi.org/10.34194/geusb.v53.8352 6 of 8 www.geusbul let in.org there is no legal requirement to do so as part of quality control at waterworks. time series analysis time series analyses could provide further insight into the way iodine concentrations vary over time. however, there is presently insufficient data coverage to conduct time series analyses across denmark; a time series analysis could be performed at only 23 of the well screens, which have at least 10 years of iodine data. all these well screens are relatively shallow (3–17.6 m b.g.l.) and located at nine grumo well sites (fig. 1), one in east denmark, the rest in west denmark. these screens are in quaternary sand aquifers except for one in a pre-quaternary sand aquifer (named ‘96.1976_2’). they are associated with 10 different groundwater bodies, as defined by troldborg (2020). the parent sediment (at 1 m depth), where known, is either proglacial (n = 4), glacial (n = 14) or postglacial (n = 1). there are mostly low iodine concentrations at these sampling points: at 18 of the well screens (78%), the median iodine concentration is ≤10 µg/l and only at three it is >20 µg/l. see table s1 for more details. figure 4 presents the time series analysis for iodine at these well screens. based on the rcv to u comparison,  the variability in iodine concentration could 141. 883_1 159. 979_2 216. 748_1 131. 1974_1 131. 1976_1 131. 1994_1 131. 1995_1 131. 1051_3 131. 1055_1 131. 1056_1 131. 1060_2 96. 1976_2 105. 1827_1 121. 958_1 131. 831_1 94. 2515_2 94. 2515_3 94. 2516_2 94. 2516_3 65. 1068_3 65. 1517_1 65. 1523_1 94. 2515_1 20 12 20 14 20 16 20 18 20 20 20 12 20 14 20 16 20 18 20 20 20 12 20 14 20 16 20 18 20 20 20 12 20 14 20 16 20 18 20 20 0 2 4 6 0 2 4 6 0 10 20 30 40 50 0 2 4 0.0 2.5 5.0 7.5 0 10 20 30 0 2 4 0 2 4 6 0 2 4 6 0.0 2.5 5.0 7.5 0 5 10 15 0 1 2 3 4 0 2 4 6 8 0 5 10 15 20 0 1 2 3 4 5 0.0 2.5 5.0 7.5 10.0 0 10 20 30 40 0.0 2.5 5.0 7.5 10.0 0 2 4 6 0 1 2 3 0.0 2.5 5.0 7.5 10.0 12.5 0.0 2.5 5.0 7.5 10.0 0 1 2 3 4 5 sampling date io di ne (μ g/ l) linear model (lm) gam model 1.7 ± 0.8 * 4.8 ± 1.3 2.9 ± 0.2 3.1 ± 0.3 5.0 ± 0.2 3.7 ± 0.2 3.9 ± 0.5 1.8 ± 0.2 10.5 ± 1.5 5.4 ± 0.7 9.7 ± 5.8 8.7 ± 1.0 3.0 ± 0.1 3.7 ± 0.2 3.4 ± 0.1 6.9 ± 0.6 * * 6.1 ± 0.5 4.8 ± 0.6 5.4 ± 0.2 2.9 ± 0.2 26.0 ± 3.0 4.7 ± 1.0 * * p < 0.05 (lm slope) median ± mad = 18.0 ± 2.0 fig. 4 iodine time series for well screens with data of at least a 10-year duration (n = 23; fig. 1; table s1 in supplementary information). the error bars show the analytical uncertainty. 95% confidence intervals are displayed for both the linear and generalised additive (gam) models. https://doi.org/10.34194/geusb.v53.8352 http://www.geusbulletin.org voutchkova 2023: geus bulletin 53. 8352. https://doi.org/10.34194/geusb.v53.8352 7 of 8 www.geusbul let in.org be attributed to environmental factors at three of these wells (’65.1068_3’, ‘105.1827_1’, ‘131.831_1’). at ’65.1068_3’, there is a positive, but not statistically significant linear trend, no non-linear trend and potentially an outlier. however, the concentrations at this well are low overall at 4.8 ± 1.3 µg/l. the well screen is shallow (4.5–5.5 m b.g.l.), located in a quaternary sand aquifer. the well screen ‘105.1827_1’ is also shallow (3–4 m b.g.l.), without a significant trend, with a possible outlier, but the iodine concentrations over the period are a bit higher at 10.5 ± 1.5 µg/l. however, at ‘131.831_1’, there is a significant decreasing trend, the screen is a bit deeper (14.5–20.5 µg/l), and the concentrations are higher, especially at the beginning of the monitoring period (9.7 ± 5.8 µg/l). two other well screens have significant linear trends (‘94.2516_3’ and ‘131.1060_2’), but both have low concentrations (3.9 ± 0.5 µg/l and 3.4 ± 0.1 µg/l, respectively) and are relatively shallow (6.5–7.5 m b.g.l. and 10.5–11.5 m b.g.l.). their variability can potentially also be attributed to analytical uncertainty. similarly, four other well screens exhibit non-linear trends (‘65.1517_1’, ‘94.2515_2’, ‘131.1051_3’, ‘216.748_1’), and their variability can also be attributed to analytical uncertainty. in addition, fig. 4 shows how the analytical uncertainty of each sample (error bars) compares to the variability, concentration level and linear/non-linear trends. the observed variability, when attributed to environmental factors, was relatively low at these well screens and so could most probably be explained by variation in the precipitation concentration and potentially by varying amounts of leaching from the soil. unfortunately, none of the locations with higher iodine concentrations had time series that were sufficiently long, and so it is presently impossible to quantify and explain the variability of the full range of iodine concentrations found in danish groundwater. conclusions this is the first time that the temporal variation of iodine in danish groundwater has been characterised systematically. the data spanned a decade (2011–2021), but the sampling frequency varied, according to the grumo programme. this nationwide assessment showed that temporal variability of groundwater iodine in denmark exceeded 20% during 2011 to 2021 at 25% of the well screens that had at least two samples (n = 930). based on comparison between the robust coefficient of variation and the analytical uncertainty, it was found that the observed variability cannot be attributed solely to analytical uncertainty for 14% of the well screens. although there were insufficient data to undertake a trend analysis at a national scale, the time series (n = 23) indicated that there could also be statistically significant linear and non-linear trends in iodine concentration at some locations. unfortunately, the time series did not include sampling locations with high iodine concentrations. moreover, the calculation of robust coefficient of variation was also limited for well screens with iodine >50 µg/l. for a quarter of the grumo well screens, there was only one sample available in the 2011–2021 period, so the variability could not be assessed at all. it is therefore recommended that the temporal variation analysis be repeated when more data have been collected as part of the grumo programme. in short, iodine in danish groundwater varies temporally; thus it may be important to evaluate this aspect in future exposure or epidemiological studies focusing on iodine in drinking water. acknowledgements i would like to thank birgitte hansen and lærke thorling for early discussions about temporal variability and the national groundwater monitoring programme (grumo). additional information funding statement the study was funded internally by the geological survey of denmark and greenland (geus). the collection of the monitoring data was funded by the danish environmental protection agency (dk epa). the conclusions and recommendations presented in this article are those of the author and do not necessarily represent the views of geus or dk epa. author contributions dv is the sole author. competing interests no competing interests. additional files three supplementary files are available with this manuscript at https:// doi.org/10.22008/fk2/ellknc 1.  aggregated_data.xlsx – data set, aggregated at well-screen-level, as used in this study. 2.  table_s1.docx – characteristics of the well screens included in the time-series analysis 3.  read_me.docx: a short description of the data set parameters in ‘aggregated_data.xlsx’. references andersen, s.b., petersen, s. & laurberg, p. 2002: iodine in drinking water in denmark is bound in humic substances. european journal of endocrinology 147, 663–670. https://doi.org/10.1530/eje.0.1470663 andersson, m., karumbunathan, v. & zimmermann, m.b. 2012: global iodine status in 2011 and trends over the past decade. the journal of nutrition 142(4), 744–750. https://doi.org/10.3945/jn.111.149393 arachchige, c.n.p.g., prendergast, l.a. & staudte, r.g. 2022: robust analogs to the coefficient of variation. journal of applied statistics 49(2), 268–290. 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https://doi.org/10.1530/eje-19-0212 wickham, h. 2016: ggplot2: elegant graphics for data analysis (v. 3.4.1). springer-verlag. https://ggplot2.tidyverse.org wood, s. 2023: mgcv: mixed gam computation vehicle with automatic smoothness estimation (r package v. 1.8-42). https://cran.r-project. org/package=mgcv who (world health organization) 2007: assessment of iodine deficiency disorders and monitoring their elimination: a guide for programme managers (3rd edition). 97 pp. geneva: world health organization. who & faoun (world health organization & food and agriculture organization of the united nations) 2004: vitamin and mineral requirements in human nutrition (2nd edition). 341 pp. geneva: world health organization. https://doi.org/10.34194/geusb.v53.8352 http://www.geusbulletin.org https://www.who.int/publications/i/item/9241592001 http://www.efsa.europa.eu/de/ndatopics/docs/ndatolerableuil.pdf http://www.efsa.europa.eu/de/ndatopics/docs/ndatolerableuil.pdf https://doi.org/10.1111/nyas.14041 https://doi.org/10.1111/nyas.14041 https://doi.org/10.17226/10026 https://doi.org/10.1016/b978-0-12-374135-6.00047-9 https://doi.org/10.1016/b978-0-12-374135-6.00047-9 https://doi.org/10.1038/s41467-022-35042-6 https://doi.org/10.1038/s41467-022-35042-6 https://cdnmedia.eurofins.com/microsites/media/1107/m050_iod_01.pdf https://cdnmedia.eurofins.com/microsites/media/1107/m050_iod_01.pdf https://www.retsinformation.dk/eli/lta/2021/2362 https://www.retsinformation.dk/eli/lta/2021/2362 https://doi.org/10.1530/eje.0.1400400 https://www.r-project.org/ https://doi.org/10.1038/sj.ejcn.1600893 https://doi.org/10.1021/acs.estlett.9b00278 https://www.geus.dk/media/8324/g02_proevetagning-okt12_uk.pdf https://www.geus.dk/media/8324/g02_proevetagning-okt12_uk.pdf https://doi.org/10.22008/gpub/32641 https://doi.org/10.22008/gpub/32641 https://doi.org/10.1016/j.scitotenv.2014.06.008 https://doi.org/10.1530/eje-19-0212 https://ggplot2.tidyverse.org https://cran.r-project.org/package=mgcv https://cran.r-project.org/package=mgcv temporal variation of iodine in danish groundwater introduction data statistical methods results and discussion nationwide temporal variability time series analysis conclusions acknowledgements additional information funding statement author contributions competing interests additional files references table table 1 recommended daily intake (rdi) and tolerable upper intake (tui) for iodine. figures fig. 1 spatial distribution of iodine sampling sites in the study period at well screens of the danish national groundwater monitoring programme (grumo). fig. 2 iodine in groundwater wells from the danish national groundwater monitoring programme (grumo). a: histogram of number of iodine samples per well screen. b: scatter plot of robust coefficient of variation (rcv) against median iodine concentration, where each symbol refers to a well screen. c: iodine concentration and its variation (expressed by the median and median absolute deviation, mad). grouped according to the number of samples per sampling location, where horizontal jitter is added to minimise overlapping points. c: grey colour is used when the mad was not calculated (only one sample available) and if mad = 0 (2–11 samples). fig. 3 iodine variability attribution at sampling sites throughout denmark. fig. 4 iodine time series for well screens with data of at least a 10-year duration (n = 23; fig. 1; table s1 in supplementary information). the error bars show the analytical uncertainty. 95% confidence intervals are displayed for both the linear and generalised additive (gam) models. research article alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 1 of 39 stratigraphy of the upper jurassic to lowermost cretaceous in the rødryggen-1 and brorson halvø-1 boreholes, wollaston forland, north-east greenland peter alsen*1 , stefan piasecki2,3,4 , henrik nøhr-hansen1, sebastian pauly5,6, emma sheldon1 , jussi hovikoski2,7 1department for geo-energy and storage, geological survey of denmark and greenland (geus), copenhagen, denmark; 2department for geophysics and sedimentary basins, geological survey of denmark and greenland (geus), copenhagen, denmark; 3globe institute, university of copenhagen, copenhagen, denmark; 4retired; 5ruhr universität bochum, bochum, germany; 6ingenum gmbh, bottrop, germany; 7geological survey of finland (gtk), espoo, finland abstract two shallow cores drilled in northern wollaston forland, north-east greenland, provide a combined section covering the upper kimmeridgian (upper jurassic) – barremian (lower cretaceous) and comprising the bernbjerg, lindemans bugt, palnatokes bjerg and stratumbjerg formations. a new lithostratigraphic unit, the storsletten member, is defined within the lindemans bugt formation. the black mudstone-dominated intervals are dated primarily by dinoflagellate cysts and ammonites, whereas the calcareous mudstones of the palnatokes bjerg formation – sandwiched between the black mudstones – are dated by calcareous nannofossils. the stratigraphy demonstrates an almost complete succession in the rødryggen-1 core, representing a deeper position in the basin, where the hiatus at the latest jurassic rift climax predicted in previous models for the eastern wollaston forland basin is absent. in contrast, the brorson halvø-1 core represents a position closer to a block crest where unconformities developed. in combination, the cores provide a key biostratigraphic reference section for the jurassic–cretaceous boundary interval in the arctic. *correspondence: pal@geus.dk received: 10 feb 2023 revised: 04 sep 2023 accepted: 05 sep 2023 published: 21 dec 2023 keywords: ammonite stratigraphy, calcareous nannofossil stratigraphy, jurassic–cretaceous boundary, palynostratigraphy, storsletten member abbreviations: aom: amorphous organic matter api: american petroleum institute fo: first occurrence geus: geological survey of denmark and greenland gr: gamma ray gssp: global boundary stratotype section and point td: total depth aff.: affinis spp.: species plural sp. juv.: species juvenile undiff.: undifferentiated geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: jon r. ineson & jørgen a. bojesen-koefoed (geus, denmark) reviewed by: simon schneider (casp, uk) and claudia schröder-adams (carleton university, canada) funding: see page 37 competing interests: see page 37 additional files: see page 37 1. introduction the rødryggen-1 and brorson halvø-1 cores were drilled in northern wollaston forland, north-east greenland (figs 1, 2). they are the second and third cores in an onshore drilling programme designed to characterise the upper jurassic source-rock succession in north-east greenland (bojesen-koefoed et al. 2014). the programme started with the drilling of the blokelv-1 core in central jameson land (fig. 1; ineson & bojesen-koefoed 2018). the rødryggen-1 borehole is located on the western side of rødryggen (meaning ‘red ridge’) at the eastern margin of storsletten, wollaston forland. previous studies in the area have suggested that the rødryggen ridge consists of upper jurassic dark mudstones, separated with a major hiatus from overlying ryazanian–hauterivian (lower cretaceous) light grey or yellowish and red mudstones and is capped by dark, mid-cretaceous mudstones (surlyk 1978; alsen 2006; pauly et al. 2013; bjerager et al. 2020; surlyk et al. 2021). coring was initiated in the valanginian and terminated in upper kimmeridgian strata at 234 m depth. the brorson halvø-1 drill site at the south-western flank of bern plateau is situated approximately 10 km to the ne from the rødryggen-1 drill site, across the sumpdalen lowland that separates the rødryggen ridge from the brorson halvø peninsula (fig. 2). coring was initiated stratigraphically somewhat higher, in barremian dark mudstones, and reached 225 m depth, also in kimmeridgian strata. https://doi.org/10.34194/geusb.v55.8342 https://orcid.org/0000-0001-6218-9054 https://orcid.org/0000-0002-7846-859x https://orcid.org/0000-0003-4353-8241 https://orcid.org/0000-0001-6330-8713 mailto:pal@geus.dk https://creativecommons.org/licenses/by/4.0/deed.ast alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 2 of 39 geusbulletin.org the blokelv-1 core in jameson land documents black mudstone deposition during the oxfordian–kimmeridgian (alsen & piasecki 2018; bjerager et al. 2018). the two cores in wollaston forland document the longevity and termination of the late jurassic black mudstone deposition in north-east greenland. the blokelv-1 core reflects a sag basin setting, whereas the wollaston forland area is characterised by block faulting and rotation and half-graben formation (surlyk 1978, 2003). the two cores drilled in wollaston forland document different depositional histories at different structural positions within the half-graben complex. hence, the brorson halvø-1 well is located near the elevated hanging wall crest of the permpas block (fig. 2; surlyk 1978), whereas the rødryggen-1 well is located near the basin centre of the same block. the aim of this study is to document in detail the lithostratigraphy and biostratigraphy of the composite cores section. this study not only adds significantly to the knowledge on the complexity of the upper jurassic stratigraphy in the area, but also presents new data of regional significance on the biostratigraphy of the jurassic–cretaceous boundary interval in the arctic. the targeted core interval, the upper jurassic dark mudstone succession, consists of two lithostratigraphic units: (1) the kimmeridgian – lower volgian bernbjerg formation and (2) a middle volgian – lower ryazanian new unit, which is here established as a new member – the storsletten member – within the lindemans bugt formation. 2. previous stratigraphic studies of the upper jurassic and lower cretaceous in wollaston forland upper jurassic – lower cretaceous strata in wollaston forland were mapped during the lauge koch–led mapping campaigns in the 1940s (vischer 1943; maync 1947). the jurassic–cretaceous boundary interval was further investigated by donovan (1964) in the northern wollaston forland, eastern kuhn ø and lindeman fjord areas (fig. 2). the jurassic – lowermost cretaceous lithostratigraphy was formally established by surlyk (1977, 1978) and recently revised and updated (surlyk et al. 2021), whereas a cretaceous lithostratigraphy was established by bjerager et al. (2020). the late jurassic ammonite faunal succession in western wollaston forland was described by sykes & surlyk (1976), whereas the ammonite and buchia bivalve zonation in the thick rift-climax succession at the jurassic–cretaceous boundary was established by surlyk (1978) and surlyk & zakharov (1982). nøhr-hansen (1993) presented a dinoflagellate cyst biostratigraphic subdivision of the barremian–albian in north-east greenland, partly based on sampled sections in wollaston forland. the zonation was recently extended and revised by nøhr-hansen et al. (2020). macrofossils and the biostratigraphy of the palnatokes bjerg formation from localities in central wollaston forland were described by alsen & rawson (2005), harper et al. (2005), alsen (2006) and alsen & mutterlose (2009), followed by calcareous nannofossil and isotope stratigraphic studies by pauly et al. (2012a) and möller et al. (2015). fig. 1 simplified geological map showing the distribution of permian– cretaceous sedimentary rocks in north-east greenland. position of study area in fig. 2 indicated with a red box. reproduced from bojesen-koefoed et al. (2023a, this volume, fig. 1). ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ g re en la nd traill ø hochstetter forland store koldewey wollaston forland jameson land cretaceous jurassic triassic permian fault inferred deep-seated fault ■ ■ miln e land kuhn ø 18°w 16°w22°w 20°w 26°w 22°w24°w28°w 76°n 72°n 26°w 24°w 100 km https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 3 of 39 geusbulletin.org fig. 2 geological map of the study area. khb: kuhn ø block. kpb: kuppel block. pb: permpas block. hb: hühnerbjerg block. modified from bojesen-koefoed et al. (2023a, this volume, fig. 1). " " major dykes/sills faults plateau lavas ice lakes rivers quaternary stratumbjerg & fosdalen fms palnatokes bjerg fm lindemans bugt fm bernbjerg fm pelion, jakobsstigen & payer dal fms permian pre-caledonian basement 10 km paleogene core-well 74°45ʹn 74°40ʹn 74°30ʹn 74°35ʹn 74°20ʹn 74°15ʹn 74°25ʹn 19°30'w 19°0ʹw 19°30ʹw 20°0ʹw20°30ʹw k u h n ø greenland zackenberg young sund ku pp el pa sse t blæ se da le n clark bjerg dronning augusta dal lindeman fjord brorson halvø kap maurer baselbjerget fligely fjord bern plateau albrecht bugt niesen lindem an sd alen clavering ø perisphinctes ravine laugeites ravine rødryggen gr æ ns er yg ge n sumpdalen w o l l a s t o n f o r l a n d khb kpb hb pb rødryggen-1 brorson halvø-1 falske bugt aucellabjerg aucellapasset st or sl et te n https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 4 of 39 geusbulletin.org 3. geological framework and lithostratigraphy the upper jurassic – lowermost cretaceous in the east greenland rift basin was deposited in a series of sub-basins between jameson land in the south at c. 70.5°n to store koldewey at c. 76.5°n in the north (fig. 1; surlyk 2003). the southernmost sub-basin in the jameson land area acted as an extensive, gently tilted platform subsiding asymmetrically with greatest subsidence to the west. north of jameson land, especially in the wollaston forland – kuhn ø area, basins are characterised by block faulting and tilting, initiated in the middle jurassic and culminating in the latest jurassic (volgian), with segmentation into narrower, strongly tilted fault blocks (surlyk 1978). middle–late jurassic deposition records an overall transgression and backstepping of the depositional system (surlyk 1977, 1978). the middle jurassic sandstone-dominated succession is thus overlain by upper jurassic offshore mudstones with a diachronous boundary, younging towards the north. the upper oxfordian – lower volgian comprises a 500–600 m thick black mudstone succession in the study area. highest sea level and maximum transgression was in the kimmeridgian, but greatest water depths were likely reached during the volgian rift climax, in the deepest parts of westerly tilted half-grabens (surlyk 2003). thick successions of very coarse, deep marine, clastic material were deposited in half-graben basins along the western, major fault scarp but rapidly passed into finer clastic deposition towards the east away from the scarp (surlyk 1978). the more easterly situated half-graben segments are almost all covered by younger successions. the late jurassic stratigraphy and deposition of those segments are thus poorly understood but are addressed by hovikoski et al. (2023a, this volume) and in the present study based on the new core data obtained from the rødryggen-1 and brorson halvø-1 cores. rifting waned in the earliest cretaceous (ryazanian–hauterivian) when half-graben sediment prisms were draped by relatively finer-grained sediments (surlyk 2003). the following early to mid-cretaceous (barremian – early albian) period was characterised by tectonic quiescence, thermal subsidence and deposition of a relatively thick succession of dark, marine siltstones and mudstones. both boreholes were initiated in the lower cretaceous and reach down into the upper jurassic. together they penetrate three major lithostratigraphic units, the vardekløft, wollaston forland and brorson halvø groups (figs 3, 4). both cores exhibit the bernbjerg, lindemans bugt and palnatokes bjerg formations. the bernbjerg formation represents the upper part of the upper jurassic (oxfordian – lower volgian) tectonostratigraphic unit, reflecting increased rifting and marine flooding (j2.4 in surlyk 2003). the lindemans bugt and palnatokes bjerg formations of the wollaston forland group represent the culmination of rifting and tilting of fault blocks (j2.5) and the end of rifting and regional drowning (j2.6), respectively. in rødryggen-1, coring began within the albrechts bugt member, the lower member of the palnatokes bjerg formation. the brorson halvø-1 core was initiated stratigraphically higher and contains the rødryggen member, the upper member of the palnatokes bjerg formation and the overlying stratumbjerg formation. the stratumbjerg formation is the lower part of an upper hauterivian – mid-albian (cretaceous) tectonostratigraphic unit defined by bjerager et al. (2020). 3.1 bernbjerg formation (vardekløft group) the upper oxfordian – lower volgian bernbjerg formation is a widespread depositional unit in north-east greenland cropping out on store koldewey in the north to traill ø in the south (surlyk et al. 2021 and references therein). its type and reference sections are located in wollaston forland and south-western kuhn ø (surlyk 1977), where the formation reaches a maximum thickness of 500–600 m (surlyk 1977; surlyk & clemmensen 1983; alsgaard et al. 2003). the bernbjerg formation sharply overlies either the payer dal formation or the jakobsstigen formation (surlyk et al. 2021). neither the rødryggen nor the brorson halvø cores reached the lower boundary of the bernbjerg formation. the nature of the upper contact to the wollaston forland group depends on location within a tilted fault block; in the downfaulted part of a fault block, the boundary is conformable, whereas the contact is an angular unconformity on the elevated fault block crests (surlyk 1977, 1978, 1991). lithologically, the formation is characterised mainly by dark grey to black mudstone and interlaminated sandstone and mudstone. sandstone beds, 5–50 cm thick, may locally show current and wave-ripple cross-stratification (surlyk 1977). surlyk (2003) referred the lowermost heterolithic unit to the ugpik ravine member. the cores described here from the rødryggen-1 and brorson halvø-1 boreholes reveal only the mud-dominated upper part of the bernbjerg formation. 3.2 lindemans bugt formation (wollaston forland group) prior to this study, middle volgian – lower ryazanian deposits were not known to crop out in the tilted permpas block, and the palnatokes bjerg formation was considered to rest directly on the bernbjerg formation. the outcropping bernbjerg formation forms a badland area east of the rødryggen ridge (figs 2, 5) towards the grænseryggen ridge. to the north, the bernbjerg formation is poorly exposed in the foot of the south-western slopes of the bern plateau (fig. 6) separated from the badland exposures in the rødryggen–grænseryggen https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 5 of 39 geusbulletin.org area by the sumpdalen valley (fig.  2). vischer (1943) and maync (1947) reported kimmeridgian black shales with ammonites, some referred to amoeboceras, others unidentified, from south of albrechts bugt and south and north of sumpdalen. the upper part of the bernbjerg formation was thus considered removed by ‘pre-valanginian’ erosion, i.e. pre-palnatokes bjerg formation, during the rift culmination at the jurassic–cretaceous transition (surlyk 1977, 1978). maync (1947) reported the kimmeridgian to be directly overlain by ‘aptian’ strata in places, i.e. the stratumbjerg formation. the poorly exposed mudstones below the albrechts bugt member at the foot of the rødryggen ridge (and at brorson halvø) were thus also considered upper jurassic bernbjerg formation, in older terminology the so-called black series (e.g. maync 1947; koch & haller 1971; surlyk 1978). however, the palynostratigraphic ages obtained in the present work from these exposed strata demonstrate the presence of middle volgian – lower ryazanian mudstones in this area, as confirmed by the borehole data. the equivalent and comparable deposits in terms of age and lithology to the west are the fine-grained laugeites ravine and, particularly, niesen members of the lindemans bugt formation. those units are, however, confined to the westernmost tilted kuppel and kuhn ø blocks (surlyk 1978) and are related to distal parts of fan deltas. we therefore consider the middle volgian – lower ryazanian mudstones in the cored sections on the permpas block as separate, detached from those other members, and thus group them into a new member, the storsletten member (fig. 3), as defined formally here (section 3.1.2.1). the unit is probably distributed throughout the permpas block, which was blocked from receiving the coarse-grained sediments that dominated deposition along the main fault to the west, since it was separated by the elevated block crest of the kuppel block and the kuhn block. other tilted block basins separated from the westernmost blocks are also likely to have been dominated by mud deposition during the middle volgian – early ryazanian. fig. 3 oxfordian–barremian lithostratigraphy for wollaston forland. l.: lower. m.: middle. u.: upper. stage formation age (ma) 125 130 135 140 145 150 155 160 stratumbjerg (pars) palnatokes bjerg lindemans bugt bernbjerg ugpik ravine red calcareous mudstone legend grey calcareous mudstone black mudstone payer dal jakobsstigen rødryggen st or sl et te n ri gi n ie se n la ug ei te s ra vi ne fa ls ke bu gt rø dr yg ge n1 b ro rs on h al vø -1 albrechts bugt young sund member sy st em se ri es lo w er ( pa rs ) c re ta c eo u s (p ar s) u pp er ju r a ss ic th et ys (s ta nd ar d) b or ea l ba rr em ia n h a u te ri v ia n va la n g ia n ki m m er id g ia n o xf o rd ia n b er ri a si a n ti th o n ia n v o lg ia n ry a z a n ia n u . u . u . l. l. u . l. l. l. l. u . u . u . m . m . w ol la st on f or la nd va rd ek lø ft b ro rs on h al vø l. su bs ta ge g ro up https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 6 of 39 geusbulletin.org fig. 4 sedimentological logs of (a) the rødryggen-1 core, (b) the brorson halvø-1 core and (c) legend to both logs. api cs.: american petroleum institute units. vf: very fine. f: fine. m: medium. c: coarse. vc: very coarse. 10 m 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 210 220 230 clay silt vf f m c f m cvc sand pebble description zo? py py py py py an an an an an an py py py py py py py p ? trace fossilsfossilssediment structures a b er nb je rg f m li nd em an s b ug t f m st or sl et te n m b w ol la st on f or la nd g ro up h al l b re dn in g g ro up pa ln at ok es b je rg f m a lb re ch ts b ug t m b (api cs.) 2000 natural gamma (api cs.) natural gamma py py py ? py py py py py py py py py py py py mfr mfr ? ? ? py py py py py py py py py py py py py py py py an an an an py show show py py dominated py py py py py py py py py py 110 120 130 140 150 160 170 180 190 200 210 220 225.7 0 m 10 20 30 40 50 60 70 80 90 100 clay silt vf cvc f m pebble an sand cmf b er nb je rg f m h al l b re dn in g g ro up li nd em an s b ug t f mw ol la st on f or la nd g ro up pa ln at ok es b je rg f m a lb re ch ts b ug t m b r ød ry gg en m b stratumbjerg formation 0 250 b https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 7 of 39 geusbulletin.org 3.2.1 storsletten member new member history. the member was not known when the lindemans bugt formation was erected (surlyk 1978). the formation encompasses mudstones deposited on the permpas block during the tectonostratigraphic phase j2.5 in surlyk (2003). type section. the rødryggen-1 core between 97 and 24.4 m depth (fig. 4a). position: n 74°32.561’, w 19°50.924’ (figs 2, 5). reference section. the brorson halvø-1 core between 45.5 and 37.5 m depth (fig. 4b). position: n 74°35.227’, w 19°34.327’ (figs 2, 6). thickness. the thickness of the member varies depending on the position on the tilted fault block. the unit is thickest where it is most complete, in the deeper, western part of the half-graben basin. it is 72.5 m thick in the rødryggen-1 core (fig. 4a). due to major hiatuses, it is only 8 m thick in the brorson halvø-1 core (fig. 4b). lithology. dark grey to black, laminated to structureless mudstone, commonly rich in pyrite and marine fossils. slump folding is common locally (brorson halvø-1 core); for lithological and diagenesis data, see hovikoski et al. (2023a, this volume) and olivarius et al. (2023, this volume), respectively. the member differs from the other coeval members by its higher content  of organic carbon and the oxygen-restricted nature of its deposits (bojesen-koefoed et al. 2023b, this volume). fossils. ammonites, buchiid bivalves, common fragments of inoceramids and palynomorphs. depositional environment. deep oxygen-restricted basin and slope, below storm wave base. the oxygen-restricted character is demonstrated by inorganic and organic geochemistry as well as the scarcity of bioturbation (bojesen-koefoed et al. 2023b, this volume; olivarius et al. 2023, this volume; hovikoski et al. 2023b). boundaries. the boundary between the bernbjerg formation and the lindemans bugt formation is lithologically transitional in the type section and occurs between 125 and 76 m (fig. 4a). the formation boundary in the type section is placed at 97 m, where the gamma-ray values increase to c. 200 api (american petroleum institute units) for the first time (bojesen-koefoed et al. 2023b, this volume). the formation change is also recognisable as gradational facies change from mudstone showing silt-clay interlamination to laminated clayey mudstone, increasing pyrite and fossil content and locally increasing slump-folding. moreover, the formation change is well-expressed in a variety of source-rock characteristics such as increasing hydrogen index, s2 and c30 desmethyl sterane values, reflecting increasing marine organic matter content (bojesen-koefoed et al. 2023b, this volume). the upper boundary towards the albrechts bugt member of the palnatokes bjerg formation is gradational in the rødryggen-1 core, whereas in the brorson halvø-1 core, the boundary is erosional and represents a hiatus (fig. 4). the gradation occurs within 1 m in the type section, mfr concretion wave ripple cross-strati�cationa contorted lamination ripple cross-strati�cation very �ne sandstone/coarse siltstone heterolithic interlamination calcite-cemented sandy mudstone ankerite and dolomite-cemented mudstone planar laminated mudstone mottled lamination pyrite ankerite slump loading synaeresis crack scour-and-�ll (gutter cast?) mud  occule ripple py an clay clast bivalve coal clasts belemnite shell fragments plant fragments chondrites burrow mottling planolites thalassinoides zoophycos ammonite onychites helminthopsis nereites mantle and swirl c fig. 4 (continued) sedimentological logs of (a) the rødryggen-1 core, (b) the brorson halvø-1 core and (c) legend to both logs. api cs.: american petroleum insti tute units. vf: very fine. f: fine. m: medium. c: coarse. vc: very coarse. https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 8 of 39 geusbulletin.org 25.5–24.4 m (fig. 4a). the boundary is readily recognisable by an abrupt decrease in gamma-ray values, increasing matrix carbonate content, a change in matrix colour from black to light grey, increasing bioturbation intensity and a change in fossil content with nannofossils, foraminifera and buchia shells becoming abundant. distribution. recorded only in the rødryggen-1 and brorson halv-1 core wells and in restricted exposures close to the well sites. chronostratigraphy. middle volgian dorsoplanites primus ammonite chronozone – lower ryazanian upper gochteodinia villosa villosa dinoflagellate cyst zone. 3.3 palnatokes bjerg formation (wollaston forland group) the upper syn-rift to early post-rift succession is represented by the palnatokes bjerg formation. it contains the albrechts bugt member, overlain by the fig. 5 views of the rødryggen-1 drill site. (a) oblique aerial view, towards the south-east of the rødryggen ridge showing the outcropping units and the position of the rødryggen-1 drill site (red dot) situated on a plateau within the yellowish weathering albrechts bugt member. the badlands area in the background is the upper jurassic bernbjerg formation. (b) view across the storsletten plain towards the east and the rødryggen ridge. ab: albrechts bugt member. pb: palnatokes bjerg formation. rr: rødryggen member. sto: storsletten member. str: stratumbjerg formation. a b sto sto ab pb (ab & rr) str str rr https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 9 of 39 geusbulletin.org rødryggen member. the young sund and falske bugt members are coarse-grained units that developed in western and eastern parts of the basin, respectively, and were not encountered in the wells drilled in the permpas block. 3.3.1 albrechts bugt member the albrechts bugt member consists of calcareous sandy, light grey to yellowish mudstone with abundant calcareous concretions (surlyk 1978; surlyk et al. 2021). at outcrop, the member weathers in conspicuous fig. 6 views of the brorson halvø-1 drill site. (a) oblique aerial view of the brorson halvø-1 drill site (red dot) at the foot of the southern slope of bern plateau, sw brorson halvø. (b) geology around the drill site. pale, yellowish weathering of the albrechts bugt member is seen sharply overlain by the rødryggen member (both palnatokes bjerg formation), overlain by the stratumbjerg formation. a roughly se–nw-oriented doleritic dyke intersects the stratumbjerg formation and a doleritic sill caps the bern plateau (sill; upper right). the drill site was situated on a small plateau in the lower part of the stratumbjerg formation (grey tent on platform). photo: a. ryge. ab: albrechts bugt member. fo: fosdalen formation. rr: rødryggen member. sto: storsletten member. str: stratumbjerg formation. a b str rr  ab sill sill fo dyke str str ab sto rr rr bern plateau fo sill dyke https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 10 of 39 geusbulletin.org bright, yellow colours that contrast strongly with the essentially grey colour of the fresh and unweathered rock. the deposits are typically highly bioturbated and rich in buchia bivalves, belemnites and ammonites (see also alsen 2006). the lithological contact between the lindemans bugt formation and albrechts bugt member of the palnatokes bjerg formation is gradational in the rødryggen-1 core and sharp in the brorson halvø-1 core. the boundary is readily recognisable by an abrupt decrease in gamma-ray (gr) values, increasing matrix carbonate and sand content, decreasing clay content, a change in matrix colour from black to light grey, increasing bioturbation intensity and a change in fossil content, with nannofossils, foraminifera and buchia shells becoming abundant. the boundary is placed at 24.4 m in the rødryggen-1 core and at 37.5 m in the brorson halvø-1 core (fig. 4). the thickness of the unit is laterally variable; it reaches a maximum thickness of c. 300 m at mt. niesen (wollaston forland), where it interfingers with the fine-grained facies of the young sund member. towards the east, on the permpas block, the member rapidly pinches out to c. 30 m (surlyk 1978). the upper boundary is transitional when overlain by the red mudstones of the rødryggen member. 3.3.2 rødryggen member (only in the brorson halvø-1 core) the rødryggen member consists of red, massive or laminated hematitic mudstones with intercalated fine sandy yellow mudstones (surlyk 1978; surlyk et al. 2021). the contact between the albrechts bugt and rødryggen members is transitional and interlayered within a 70 cm interval in the brorson halvø-1 core. the change is visible as a gradational change in colour from light grey to red; the boundary is placed at 30.0 m where the matrix turns permanently red (fig. 4b). moreover, the rødryggen member differs from the albrechts bugt member in microfossil and macrofossil content with foraminifers and inoceramid shell becoming more common. 3.4 stratumbjerg formation (brorson halvø group; only in brorson halvø-1 core) the stratumbjerg formation marks a return to deposition of dark, fine-grained sediments in an oxygen-restricted environment below storm wave base during the tectonically quiescent phase after the volgian rift climax and ryazanian–valanginian late rift phases. the unit is widely distributed throughout north-east greenland, from traill ø in the south to store koldewey in the north (bjerager et al. 2020). the boundary to the underlying rødryggen member is gradational within an 80 cm interval in the brorson halvø-1 core. the boundary is placed at 8.7 m (top of transition), where the bioturbated grey mudstones no longer interfinger with reddish mudstones (fig. 4b). 4. biostratigraphic methods and approach the rødryggen-1 borehole is 234.5 m deep, and the brorson halvø-1 borehole is 225.7 m deep. both boreholes were fully cored, with recoveries of 99%. the core diameter in both cores is 42 mm. each core is essentially treated as one sample and assigned codes with the prefix ‘geus’. the rødryggen-1 well is designated as geus 517001, and the brorson halvø-1 is designated as geus 517003. sub-numbers were assigned to the extracted material, which was sampled for a suite of core analysis, including biostratigraphy, geochemistry, rock properties and provenance (see also bojesen-koefoed et al. 2023b, this volume; olivarius et al. 2023, this volume). for simplicity, in this study, we only refer to the levels or borehole depths (below surface) for the respective subsamples, for example in range charts, tables and figure captions. 4.1 macrofossils before the cores were subject to slabbing and sampling for the standard analytical programme, they were examined for macrofossils. examination for macrofossils was particularly directed to where the core had naturally split along bedding planes, commonly along planes with fossils, so that all ends of the individual core pieces were inspected. most of the fossils are fragmented. identification also suffers from the relatively small diameter of the core such that only small portions of a fossil were usually available for study, and thus fewer diagnostic characters are available in these cases. 4.2 palynomorphs and calcareous nannofossils mudstones for palynostratigraphic analysis were initially sampled at an even spacing throughout the core. upon initial biostratigraphic screening, additional material was subsequently sampled in selected intervals, for example, across lithostratigraphic or chronostratigraphic boundaries such as the jurassic–cretaceous boundary interval, to obtain higher biostratigraphic subdivision and precision. the calcareous mudstone intervals of the palnatokes bjerg formation were sampled specifically for calcareous nannofossil stratigraphy. each sample for palynomorphs and nannofossil analysis comprised a split/slabbed, 4 to 6 cm thick core interval. the well depth of a mudstone sample is the medium point of the thickness of the sample. the palynological preparation methods of the crushed sample material include processing with hcl, https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 11 of 39 geusbulletin.org hf, oxidation with hno3 and heavy-liquid separation. samples from the bernbjerg and lindemans bugt formations contain abundant marine and terrestrial organic matter, which required repeated oxidation and ultrasonic treatment to release identifiable dinoflagellate cysts, with slides produced after each repeated preparation step. when a state of sufficient oxidation was reached, the organic residue was sieved with a 21 µm filter. the coarser fraction was swirled and eventually mounted on glass slides using a glycerine jelly medium. the dinoflagellate cyst content was analysed using a normal light microscope. all dinoflagellate cysts in one slide from each sample were counted to perform a semiquantitative analysis. this approach applies to dark mudstone intervals, whereas the diversity and abundance in the very low organic, calcareous mudstones of the palnatokes bjerg formation do not allow for comparison with data from the dark mudstone units. in addition to dinoflagellate cysts, acritarchs and phrasinophycean and freshwater algae were also counted. the palynological taxonomy follows the lentin and williams index of fossil dinoflagellates, 2004 edition (fensome & williams 2004), unless otherwise indicated by author references. nannofossil slides were prepared using the simple smear slide technique of bown & young (1998). where samples appeared to be barren with respect to calcareous nannofossils, three length traverses of the smear slide were examined. where a particular species of nannofossil dominated the slide, one length traverse was counted, and then two further lengths were checked for rare forms. biostratigraphic ranges of nannofossils are adapted from burnett (1998), bown et al. (1998) and pauly et al. (2012a). 4.3 stratigraphic nomenclature and methodology the tithonian and berriasian are the standard stages for the uppermost jurassic and lowermost cretaceous, respectively, as defined in the tethyan realm. due to pronounced faunal provincialism around the system boundary and the recognition of faunally clearly separated tethyan and boreal realms, a parallel stage nomenclature has evolved for the boreal area. until improved correlation between boreal and tethyan areas is obtained, a boreal subdivision and stage nomenclature are commonly adopted in greenland, including in this study. this study also follows the stratigraphic concept commonly applied for middle–upper jurassic stratigraphic studies in east greenland, which considers ammonite zones as chronozones, representing rock units that are also identifiable by means of fossil groups other than ammonites. the east greenland upper jurassic ammonite zonation is thus closely integrated with the dinoflagellate cyst record (see also discussion in alsen & piasecki 2018 and references therein). 4.4 rødryggen-1 core the rødryggen-1 core was initially sampled at 37 levels for palynostratigraphy and nannofossil stratigraphy with a standard c. 10 m spacing, but with somewhat denser sampling in the lowermost and uppermost parts of the core. subsequent sampling at critical levels resulted in a total of 63 sampled levels being analysed for palynostratigraphy. the dark mudstones of the bernbjerg, lindemans bugt and stratumbjerg formations were primarily dated with palynostratigraphy aided by ammonite stratigraphy. the pale, calcareous mudstones of the albrechts bugt member (palnatokes bjerg formation) were dated by a combination of palynostratigraphy and calcareous nannofossil stratigraphy. ammonites occur only in two intervals, between 225 and 197 m, and in a c. 65 m thick interval between 90 and 26 m (fig. 4a). the few ammonites in the lower interval are of little biostratigraphic value (table 1), so age-significant ammonites are essentially restricted to the storsletten member of the lindemans bugt formation. no ammonites were found in the interval between 90 and roughly 200 m. the latter interval is characterised by finely laminated dark mudstones, which theoretically have a good preservation potential for fossils. ammonites are thus probably present. however, the interval is intensely fractured hampering both the preservation and detection of fossils. ten samples from the rødryggen-1 core were examined for nannofossil content. seven mudstone samples from the bernbjerg and lindemans bugt formations were almost barren with respect to calcareous nannofossils. thus, only the three samples from the calcareous mudstones of the albrechts bugt member (palnatokes bjerg formation) yielded relatively good nannofossil recovery. 4.5 brorson halvø-1 core the brorson halvø-1 core was sampled for palynostratigraphy and nannofossil stratigraphy with default sample spacing of c. 10–15 m, a slightly less dense sampling strategy than used for the rødryggen-1 core. subsequent closer sampling was undertaken in lithostratigraphic and chronostratigraphic boundary intervals to obtain higher precision, for example, to ascertain the stratigraphic significance of hiatuses. well-preserved ammonites are absent in the core and the integration of ammonite and palynostratigraphy, which was of great benefit in the analysis of the rødryggen-1 core, could not be undertaken. the biostratigraphy of the core receives only minor support from the https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 12 of 39 geusbulletin.org table 1 summary of macrofossils recorded in the rødryggen-1 core. depth geus 517001 and sub-sample number mguh no.a figure no. (this study) description / taxononomy faunal horizon sensu callomon & birkelund (1982) ammonite stratigraphy 27.73 m –401 34200 10.s hectoroceras sp.   h. kochi zone 35.80 m –402     bivalve   p. maynci zone 37.49 m –438 34199 10.r cf. praetollia maynci spath   p. maynci zone 40.22 m –437     buchia bivalve   s. (sw.) primitivus zone 43.18 m not sampled     ammonite indet.   s. (sw.) primitivus zone 43.21 m –403 34198 10.q subcraspedites (swinnertonia) sp. juv.   s. (sw.) primitivus zone 45.37 m –404 34197 10.p s. (swinnertonia) cf. subundulatus swinnerton   s. (sw.) primitivus zone 45.73 m –405     bivalve   (p. exoticus zone) 47.20 m –406 34196 10.o aptychus   (p. exoticus zone) 47.52 m –407     bivalve   (p. exoticus zone) 47.94 m –436     ?praechetaites sp.   (p. exoticus zone) 48.09 m –435     buchia bivalve   (p. exoticus zone) 48.37 m –408 34195 10.n ammonite indet.   (p. exoticus zone) 48.54 m –427 34194 10.m cf. praechetaites exoticus (shulgina)   (p. exoticus zone) 48.82 m –434 34193 10.l cf. praechetaites exoticus (shulgina)   (p. exoticus zone) 49.94 m –433     ammonite; indeterminate juvenile   (p. exoticus zone) 50.14 m –432     cf. praechetaites exoticus (shulgina)   (p. exoticus zone) 51.55 m –409 34192 10.k laugeites cf. planus mesezhnikov m 47 l. groenlandicus zone 52.30 m –431     ammonite fragment; indeterminate   l. groenlandicus zone 52.50 m –430     l. cf. biplicatus m 47 l. groenlandicus zone 52.82 m –410 34191 10.j laugeites cf. intermedium donovan   l. groenlandicus zone 53.27 m –429     ammonite aptychi   l. groenlandicus zone 53.67 m –411 34190 10.i laugeites cf. biplicatus mesezhnikov m 47 l. groenlandicus zone 54.53 m –412     ?laugeites sp.   l. groenlandicus zone 55.25 m –413     ammonite indet.   e. pseudapertum zone 55.98 m –414 34189 10.h epipalliceras cf. pseudapertum spath m 42 e. pseudapertum zone 62.43 m –415 34188 15.b micro-onychites   d. gracilis zone 66.54 m –416     belemnite   d. gracilis zone 66.84 m –417     ammonite indet.   d. gracilis zone 70.06 m –418 34187 10.g dorsoplanites jamesoni spath m 40? d. gracilis zone 74.03 m –419 34186 10.f pavlovia cf. corona callomon & birkelund m 37 d. liostracus zone 74.20 m –420 34185 10.e dorsoplanites aff. liostracus m 37 d. liostracus zone 75.43 m –421 34184 10.d pavlovia cf. variocostata callomon & birkelund m 35 p. communis zone 89.26 m –422 34183 15.a mega-onychites   d. primus zone 89.30 m –423 34182 10.c dorsoplanites primus callomon & birkelund m 31 d. primus zone 90.85 m –424 34181 10.b dorsoplanites sp.   d. primus zone 199.42 m –425     ammonite indet.   ? 227.72 m –426 34180 10.a amoeboceras? sp.   ? athe specimens are stored in the palaeontology type collection at the natural history museum of denmark and each labelled with an mguh number – museum geologica universitas hafniensis.   https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 13 of 39 geusbulletin.org macrofossils collected in exposures near the drill site, where alsen (2006) recorded the ammonite fauna in the albrechts bugt and rødryggen members. a total of 41 samples were analysed for palynomorphs: 19 in the bernbjerg formation, 7 in the lindemans bugt formation (storsletten member), 12 in the palnatokes bjerg formation and 3 in the stratumbjerg formation. from the upper part of the brorson halvø-1 core (40.27 to 6.16 m), analysis of calcareous nannoplankton was applied in 19 samples. the lowest sample represents dark mudstone of the lindemans bugt formation and is barren, as are equivalent black mudstone samples from the rødryggen-1 core. in the remaining samples, the preservation of the calcareous nannoplankton varies from good to moderate. most samples of the albrechts bugt member yield well-preserved calcareous nannoplankton. the calcareous nannoplankton in the rødryggen member shows good to moderate preservation and overgrowth by iron oxide minerals. surprisingly, the organic-rich grey mudstones of the stratumbjerg formation also yielded well-preserved calcareous nannoplankton assemblages. 5. biostratigraphy of the rødryggen-1 core the biostratigraphic subdivision of the rødryggen-1 core is described from total depth (td) at 234.40 m upwards. the location of bulk-rock samples for palynostratigraphy is seen in the charts illustrating the distribution and ranges of dinoflagellate cyst taxa, palyno events and calcareous nannofossil taxa (figs 7, 8). the recorded ammonite levels are listed in table 1 and illustrated in a stratigraphic distribution chart (fig. 9). selected ammonites and dinoflagellate cysts are illustrated in figs 10–13. 5.1 aulacostephanus eudoxus chronozone (234.40 m (td) – 220.51 m) fossils. the recognition of the zone is based on its dinoflagellate cyst record. one poorly preserved ammonite is a possible amoeboceras (227.72 m; fig. 10a; table 1). the base of the zone is arbitrarily placed at the base of the core (at td). the dinoflagellate cyst assemblage is relatively diverse, and cysts are abundant. the zone is dominated by perisseiasphaeridium pannosum, paragonyaulacysta capillosa and cribroperidinium spp., whereas epiplosphaera reticulospinosa and paragonyaulacysta borealis are common taxa at discrete levels within the zone. biostratigraphy. abundant p. pannosum is reported to range from ammonite faunal horizon m 20, at the base of the aulacostephanus eudoxus zone, to a level between faunal horizons m 22 and m 23 near the top of the a. eudoxus zone in milne land (fig. 14; piasecki 1996). p. pannosum was used as a key taxon of the a. eudoxus zone in the blokelv-1 core in jameson land (alsen & piasecki 2018) and the brorson halvø-1 core (herein). rhycodiniopsis cladophora has its uppermost occurrence at m 21 in milne land (piasecki 1996). p. borealis is common in these assemblages and is present from 234.41 m, but it is known to range much deeper, to the p. baylei zone level in milne land (piasecki 1996). dingodinium minutum, gonyaulacysta jurassica, taeniophora iunctispina and possibly nannoceratopsis pellucida occur scattered and are rare in this chronozone, either as last occurrences or reworked from strata below. age. kimmeridgian, late jurassic. organic matter. amorphous kerogen (often termed aom) together with a large proportion of terrestrial organic material from higher land plants, especially degraded black grains of woody material, dominate the interval. the presence of dinoflagellate cysts indicates a marine environment, and the dominance of aom is suggestive of oxygen-deficient bottom conditions. 5.2 aulacostephanus autissiodorensis chronozone (220.51–150.25 m) fossils. the recognition of this chronozone is based on its dinoflagellate cyst record. it contains only one indeterminable ammonite (199.42 m; table 1). the base of the zone is recognised as the level with the lowest occurrence of abundant oligosphaeridium patulum (220.51 m). the composition of the dinoflagellate cyst assemblage changes from 220.51 m to be characterised by abundant oligosphaeridium patulum, cribroperidinium spp. and cyclonephelium distinctum. paragonyaulacysta capillosa is common but less frequent upwards. the assemblage is characterised by low diversity and low abundance, probably due to the high content of organic material. biostratigraphy. abundant o. patulum appears in the basal a. autissiodorensis chronozone between ammonite faunal horizons m 22 and m 23 in milne land (fig. 14; piasecki 1996; alsen & piasecki 2018). cribroperidinium complexum has its highest occurrence at the boundary of the a. autissiodorensis and p. elegans chronozones (biostrat 2018). age. kimmeridgian, late jurassic. organic matter. the interval is characterised by abundant amorphous kerogen together with a large proportion of terrestrial organic material from higher land plants, especially brown to black woody material. the https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 14 of 39 geusbulletin.org fi g. 7 t he r an ge s of k ey d in ofl ag el la te c ys t t ax a in th e rø dr yg ge n1 co re . t he d in ofl ag el la te c ys t s pe ci es a re s tr at ig ra ph ic al ly a rr an ge d ac co rd in g to th e su cc es si on o f t he ir lo w es t o cc ur re nc es . a r an ge c ha rt s ho w in g al l ta xa r ec or de d in th e co re is a va ila bl e in s up pl em en ta ry d at a fi le 1 . 1. 10 6. 59 9. 00 10 .1 7 15 .8 5 17 .8 2 to p of l ag en or hy tis d el ic at ul a to p of g oc ht eo di ni a vi llo sa v ill os a to p of s ys te m at op ho ra p al m ul a b as e of n el ch in op sis k os tr om ie ns is to p of t ub ot ub er el la a pa te la b as e of a bu nd an t o lig os ph ae ri di um c om pl ex 20 .4 4 24 .1 0 25 .8 2 27 .0 0 29 .8 9 31 .9 5 35 .0 2 37 .7 1 to p of p ar ag on ya ul ac ys ta c ap ill os a b as e of p ho be ro cy st a ne oc om ic a to p of p ar ag on ya ul ac ys ta b or ea lis to p of s cr in io di ni um p ha ro to p of r ot os ph ae ro ps is th ul a b as e of p al ae cy st a pa lm ul a b as e of s cr in io di ni um p ha ro b as e of c an ni ng ia c om pt a b as e of is th m oc ys tis d is tin ct a b as e of c om m on g oc ht eo di ni a vi llo sa v ill os a 38 .2 8 to p of a bu nd an t p te ro sp er m el la s pp . 40 .6 2 b as e of a bu nd an t p te ro sp er m el la s pp . 56 .8 8 59 .8 5 to p of l ep to di ni um su bt ile 70 .1 5 b as e of l ag en or hy tis d el ic at ul a 79 .7 4 b as e of m ud er on gi a sim pl ex 90 .2 7 a bu nd an t c as si cu lo sp ha er id ia s pp . 99 .7 2 to p of o lig os ph ae ri di um p at ul um 11 0. 15 to p of t ri ch od in iu m p ia se ck ii to p of s en on ia sp ha er a cl av el li 13 0. 26 to p of a bu nd an t o lig os ph ae ri di um pa tu lu m 14 0. 25 b as e of s en on ia sp ha er a cl av el li 15 0. 25 b as e of t ri ch od in iu m p ia se ck ii to p of c ri br op er id in iu m c om pl ex um 22 0. 51 to p of p er is se ia sp ha er id iu m p an no su m b as e of o lig os ph ae ri di um p at ul um 22 8. 94 to p of g on ya ul ac ys ta ju ra ss ic a boreal early cretaceous boreal late jurassicperiod/epoch albrechts bugt storslettenmember palnatokes bjerg lindemans bugtformation bernbjerg wollaston forlandgroup hall bredning 5 10 15 20 25 30 35 40 45 50 55 60 65 70 75 80 85 90 95 1 00 1 05 1 10 1 15 1 20 1 25 1 30 1 35 1 40 1 45 1 50 1 55 1 60 1 65 1 70 1 75 1 80 1 85 1 90 1 95 2 00 2 05 2 10 2 15 2 20 2 25 2 30 td depth (m) li th os tr at ig ra ph y l. v ol gi an m. volgian e. volgian kimmeridgian la te v al an gi ni an ea rl y v al an gi ni an la te r ya za ni an ea rl y r ya za ni an age up pe r o . c om pl ex lo w er o . c om pl ex up pe r g . v ill os a vi llo sa lo w er g . v ill os a vi llo sa zone zone chronostratigraphy dino�agellate cyst zones (nøhr-hansen et al. 2020) ammonite chronozones (callomon & birkelund 1982) a .e ud ox us p. e le ga ns a. autissiodorensis p. hu dd le st on e p. p ec tin at us d . p ri m us (p. ia tr ie ns is p. ru go sa ) p. c om m un is d . l io st ra cu s e. p se ud ap er tu m l. g ro en la nd ic us p. e xo tic us s. p ri m iti vu s p. m ay nc i h . k oc hi d . g ra ci lis (p . s ci tu lu s ) p. w he at le ye ns is perisseiasphaeridium pannosum d in o� ag el la te c ys ts a cr it ar ch s cribroperidinium complexum oligosphaeridium patulum nelchinopsis kostromiensis oligosphaeridium complex valensiella phoberocysta neocomica rotosphaeropsis thule gochteodinia villosa villosa perisseiasphaeridium insolitum lagenorhytis delicatula muderongia simplex ev en ts st ra ti gr ap hi c ra ng e st ra ti gr ap hi c ra ng e r leptodinium subtile r r rgonyaulacysta jurassica r paragonyaulacysta capillosa + senoniasphaera clavelli trichodinium piaseckii paragonyaulacysta borealis tubotuberella apatela palaecysta palmula pterospermella spp. pterospermella sp. (small) scriniodinium pharo canningia compta isthmocystis distincta cassiculosphaeridia magna g oc ht eo di ni a vi llo sa v ill os a b as e of p er is se ia sp ha er id iu m in so lit um 1. 10 6. 59 9. 00 10 .1 7 15 .8 5 17 .8 2 to p of l ag en or hy tis d el ic at ul a to p of g oc ht eo di ni a vi llo sa v ill os a to p of s ys te m at op ho ra p al m ul a b as e of n el ch in op sis k os tr om ie ns is to p of t ub ot ub er el la a pa te la b as e of a bu nd an t o lig os ph ae ri di um c om pl ex 20 .4 4 24 .1 0 25 .8 2 27 .0 0 29 .8 9 31 .9 5 35 .0 2 37 .7 1 to p of p ar ag on ya ul ac ys ta c ap ill os a b as e of p ho be ro cy st a ne oc om ic a to p of p ar ag on ya ul ac ys ta b or ea lis to p of s cr in io di ni um p ha ro to p of r ot os ph ae ro ps is th ul a b as e of p al ae cy st a pa lm ul a b as e of s cr in io di ni um p ha ro b as e of c an ni ng ia c om pt a b as e of is th m oc ys tis d is tin ct a b as e of c om m on g oc ht eo di ni a vi llo sa v ill os a 38 .2 8 to p of a bu nd an t p te ro sp er m el la s pp . 40 .6 2 b as e of a bu nd an t p te ro sp er m el la s pp . https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 15 of 39 geusbulletin.org fig. 8 the distribution of calcareous nannofossils in the rødryggen-1 core. d ep th li th os tr at ig ra ph y 3.60 pa ln at ok es bj er g a lb re ch ts bu gt lo w er c re ta ce ou s early valanginian late ryazanian e ryazanian24.40 fo rm at io n 3.60 24.40 m em be r c hr on os tr at ig ra ph y 3.60 pe rio d/ ep oc h 6.59 15.85 a ge lo w er c re ta ce ou s na nn of os si le s (b ow n et a l. 19 98 ) bc3b zo ne sa m pl es 3.60m 13.82m 16.60m semi-quantitative nanno abundance scheme c al cu lit es ? s p. 1 n an no co nu s c on ca vu s rh ag od is cu s a sp er st au ro lit hi te s s tr ad ne ri th or ac os ph ae ra s pp . tr an ol ith us g ab al us w at zn au er ia b ar ne sia e w at zn au er ia b rit an ni ca w at zn au er ia f os sa ci nc ta a ss ip et ra in fr ac re ta ce a c ru ci bi sc ut um sa le br os um c yc la ge lo sp ha er a m ar ge re lii d ia zo m at ol ith us le hm an ii re te ca ps a an gu st ifo ra ta re te ca ps a cr en ul at a ro te la pi llu s c re nu la tu s st ra dn er lit hu s s ilv ar ad iu s te gu m en tu m st ra dn er i tr ip in na lit hu s s he tla nd en si s w at zn au er ia m an iv ita e w at zn au er ia o va ta ze ug rh ab do tu s e m be rg er i ze ug rh ab do tu s e re ct us ze ug rh ab do tu s � ss us n an no co nu s s pp . re te ca ps a sc hi zo br ac hi at a so lla sit es h or tic us ? nannopalaeontology nanno abundance scheme rare (1) few (5) common (50) abundant (75) dominant (100+) 20 m fig. 9 the ranges of ammonite taxa recorded in the rødryggen-1 core. d ep th (m ) lithostratigraphy g ro up fo rm at io n m em be r c hr on os tr at ig ra ph y pe ri od /e po ch a ge a m m on it e zo na ti on (c al lo m on & b ir ke lu nd 1 98 2) zo ne samples sa m pl e de pt h is b a se o f d ep th ra ng e 27.73m 35.80m 37.49m 40.22m 43.18m 43.21m 45.37m 45.73m 47.20m 47.52m 47.94m 48.09m 48.37m 48.54m 48.82m 49.94m 50.14m 51.55m 52.30m 52.50m 52.82m 53.23m 53.27m 53.67m 54.53m 55.25m 55.98m 62.43m 63.84m 66.54m 66.84m 70.06m 74.03m 74.20m 75.43m 89.26m 89.30m 90.85m pa ln at ok es b je rg fo rm at io n li nd em an s b ug t f or m at io n a lb re ch ts b ug t m em be r b or ea l e ar ly c re ta ce ou s early ryazanian early valanginian late ryazanian ammonites 5 10 15 20 25 30 35 40 45 50 55 60 65 70 75 80 85 90 95 d or so pl an ite s s p. d or so pl an ite s p ri m us c al lo m on & b ir ke lu nd m eg aon yc hi te s pa vl ov ia v ar io co st at a c al lo m on & b ir ke lu nd d or so pl an ite s a �. li os tr ac us pa vl ov ia c f. co ro na c al lo m on & b ir ke lu nd d or so pl an ite s g ra ci lis s pa th m ic ro -o ny ch it es ep ip al lic er as c f. ps eu da pe rt um s pa th la ug ei te s c f. bi pl ic at us m es ez hn ik ov la ug ei te s b ip lic at us m es ez hn ik ov a pt yc hu s s p. la ug ei te s c f. in te rm ed iu m d on ov an la ug ei te s c f. pl an us m es ez hn ik ov pr ae ch et ai te s c f. ex ot ic us ( sh ul gi na ) pr ae ch et ai te s s pp . su bc ra sp ed ite s ( sw in ne rt on ia ) s ub un du la tu s s w in ne rt on su bc ra sp ed ite s ( sw in ne rt on ia ) s p. ju v. pr ae to lli a cf . m ay nc i s pa th h ec to ro ce ra s s p. stratigraphic range w ol la st on f or la nd g ro up st or sl et te n m em be r b or ea l l at e ju ra ss ic m id dl e v ol gi an late volgian late valanginian d. primus (-p. iatriensis p. rugosa) p. communis d. liostracus e. pseudapertum l. groenlandicus p. exoticus s. primitivus p. maynci h. kochi d. gracilis https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 16 of 39 geusbulletin.org fig. 10 selected ammonites recorded in the rødryggen-1 core. a: amoeboceras? sp., level 227.72 m, mguh 34180. b: dorsoplanites sp., level 90.85 m, mguh 34181. c: dorsoplanites primus, level 89.30 m, mguh 34182. d: pavlovia cf. variocostata, level 75.43 m, mguh 34184. e: dorsoplanites aff. liostracus, level 74.20 m, mguh 34185. f: pavlovia cf. corona, level 74.03 m, mguh 34186. g: dorsoplanites jamesoni, level 70.06 m, mguh 34187. h: epipallasiceras cf. pseudapertum, level 55.98 m, mguh 34189. i: laugeites cf. biplicatus, level 53.67 m, mguh 34190. j: laugeites cf. intermedium, level 52.82 m, mguh 34191. k: laugeites cf. planus, level 51.55 m, mguh 34192. l: cf. praechetaites exoticus, level 48.82 m, mguh 34193. m: cf. praechetaites exoticus, level 48.54 m, mguh 34194. n: ammonoidea indet., level 48.37 m, mguh 34195. o: aptychus, level 47.20 m, mguh 34196. p: s. (swinnertonia) cf. subundulatus, level 45.37 m, mguh 34197. q: subcraspedites (swinnertonia) sp. juv., level. 43.21 m, mguh 34198. r: cf. praetollia maynci spath, level 37.49 m, mguh 34199. s: hectoroceras sp., level 27.73 m, mguh 34200. the specimens are stored in the palaeontology type collection at the natural history museum of denmark and each labelled with an mguh number – museum geologica universitas hafniensis. a g j l m n b c d e f h i k o p sq 1 2 3cm r https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 17 of 39 geusbulletin.org a b c d e f g h i j k l m n o p fig. 11 selected biostratigraphically significant dinoflagellate cysts and acritarchs from the rødryggen-1 core. scale bars: 25 µm. a and b: paragonyaulacysta capillosa, sample 234.4 m, slide 2. c and d: paragonyaulacysta borealis, sample 140.25 m, slide 6, and 136.26 m, slide 5. e and f: perisseiasphaeridium pannosum, sample 226.22 m, slide 4. g and h: oligosphaeridium patulum, sample 130.26 m, slide 6. i and j: cribroperidinium complexum, low and high focus on the same specimen, sample 234.5 m, slide 5. k and l: trichodinium piaseckii, sample 136.26 m, slide 6. m: senoniasphaera clavellii, sample 90.27 m, slide 5. n: cassiculosphaeridium magna, sample 90.27 m, slide 5. o: wallodinium krutzschii, sample 110.15 m, slide 4. p: muderongia simplex, sample 76.74 m, slide 4. https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 18 of 39 geusbulletin.org fig. 12 selected biostratigraphically significant dinoflagellate cysts and acritarchs from the rødryggen-1 core. scale bars: 25 µm. a, b and c: lagenorhytis delicatula, showing variable morphology, sample 54.31 m, slide 7. d: gochteodinia villosa subsp. villosa, sample 56.88 m, slide 4 and 5. e: leptodinium subtile, sample 90.27 m, slide 5. f: perisseiasphaeridium insolitum, sample 59.85 m, slide 4. g, h and i: large pterospermella spp., sample 40.62 m, slide 5, and sample 35.2 m, slide 7. j and k: small pterospermella spp., sample 43.43 m, slide 5 and sample 35.2, slide 7. l and m: isthmocystis distincta, sample 35.02 m, slide 3. n, o, p, q and r: morphological variations of gochteodinia villosa subsp. villosa, (n–q) sample 37.71 m, slide 7, and (r) sample 35.02 m, slide 4. s: circulodinium compta, sample 31.95 m, slide 4. t: scriniodinium pharo, sample 24.1 m, slide 3. a b c d e f g h i j k l m n o p q r s t https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 19 of 39 geusbulletin.org fig. 13 selected biostratigraphically significant dinoflagellate cysts and acritarchs from the rødryggen-1 (a–m) and the brorson halvø-1 (n–r) cores. scale bars: 25 µm. a: palaecysta palmula, sample 24.1 m, slide 2. b: rotosphaeropsis thule, sample 27.0 m, slide 7. c: scriniodinium pharo, sample 27.1 m, slide 3. d: paragonyaulacysta borealis, sample 35.2 m, slide 4. e: tubotuberella apatela, sample 25.0 m, slide 9. f: phoberocysta neocomica, sample 13.2 m, slide 2. g: paragonyaulacysta capillosa, sample 21.38 m, slide 2. h: oligosphaeridium complex, sample 10.17 m, slide 2. i: oligosphaeridium complex, sample 17.82 m, slide 2. j: nelchinopsis kostromiensis, sample 10.17 m, slide 2. k: palaecysta palmula, sample 9.00 m, slide 2. l: gochteodinia villosa subsp. villosa, sample 24.1 m, slide 3. m: lagenorhytis delicatula, sample 2.78 m, slide 4. n: batioladinium longicornutum, sample 3.37 m, slide 3. o: muderongia tetracantha, sample 9.18 m, slide 3. p: muderongia staurota, sample 9.18 m, slide 3. q: nelchinopsis kostromiensis, sample 6.16 m, slide 3. r: pseudoceratium anaphrissum, sample 3.37 m, slide 3. a b c d e f g h i j k l m n o p q r https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 20 of 39 geusbulletin.org praechetaites tenuicostatus zone pectinatites wheatleyensis zone praetollia maynci zone hectoroceras kochi zone pictonia baylei zone dorsoplanites gracilis zone crendonites anguinus zone aulacostephanus mutabilis zone epipallasiceras pseudapertum zone dorsoplanites liostracus zone dorsoplanites primus zone pectinatites (virgatosphinctes) elegans zone pavlovia rugosa zone pavlovia iatriensis zone p. eastlecottensis subzone aulacostephanus autissiodorensis zone paracrioceras elegans zone aulacostephanus eudoxus zone peregrinus albidum zone pectinatites pectinatus zone pectinatites hudlestoni zone p. paravirgatus subzone pavlovia communis zone rasenia cymodoce zone pectinatites (virgatosphinctes) scitulus zone polyptychites michalskii zone dichotomites hollwedensis zone laugeites groenlandicus zone praechetaites exoticus zone dichotomites crassus zone dichotomites bidichotomoides zone subcraspedites primitivus zone simbirskites decheni zone simbirskites (speetoniceras) inversum zone fissicostaciceras  ssicostatum zone surites analogus zone surites tzikwinianus zone delphinites undulatoplicatilis zone nikitinoceras hoplitoides zone paracrioceras denckmanni zone parancyloceras bidentatum zone biostratigraphic hiatus biostratigraphic hiatus biostratigraphic hiatus biostratigraphic hiatus biostratigraphic hiatus biostratigraphic hiatus biostratigraphic hiatus epilaugeites surlyki zone (vogulicus zone sensu surlyk 1978) c re ta c eo u s (p ar s) ju ra ss ic lo w er (p ar s) u pp er ba rr em ia n h a u te ri v ia n va la n g ia n ki m m er id g ia n v o lg ia n ry a za n ia n u . u . u . l. l. u . u . m . l. l. u . l. l. chronostratigraphy zone/subzone m 47 m 46 m 45 m 44 m 43 m 42 m 41 m 40 m 39 m 38 m 37 m 36 m 35 m 34 m 33 m 32 m 31 m 30 m 29 m 28 m 27 m 26 m 24 m 23 m 22 m 21 m 20 m 19 m 18 m 17 m 16 m 15 m 14 m 25 fig. 14 kimmeridgian – barremian ammonite zonation for north-east greenland. m 14 to m 47 are faunal horizons recorded in milne land by callomon & birkelund (1982). l.: lower. m.: middle. u.: upper. https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 21 of 39 geusbulletin.org presence of dinoflagellate cysts indicates a marine environment, and the dominance of aom is suggestive of oxygen-deficient bottom conditions. 5.3 pectinatites elegans chronozone (150.25–130.26 m) fossils. the recognition of this chronozone is based on its dinoflagellate cyst record. the dinoflagellate cyst assemblage is characterised by abundant oligosphaeridium patulum, cribroperidinium spp., sirmiodinium grossii and c. distinctum. paragonyaulacysta capillosa is common but less frequent upwards. the assemblage is mostly of low diversity and low abundance, probably due to the abundance of organic material. biostratigraphy. the base of the zone is recognised as the level with the last occurrence of cribroperidinium complexum and the first occurrence of trichodinium piaseckii. age. early volgian, late jurassic organic matter. the interval is characterised by abundant amorphous kerogen together with a large proportion of terrestrial organic material from higher land plants, especially brown to black woody material. the presence of dinoflagellate cysts indicates a marine environment, and the dominance of aom is suggestive of oxygen-deficient bottom conditions. 5.4 (pectinatites scitulus –) pectinatites wheatleyensis chronozones undiff. (130.26–119.70 m) fossils. the recognition of the interval is based on its dinoflagellate cyst record. the base is placed above the last occurrence of abundant oligosphaeridium patulum. the dinoflagellate cyst assemblage is characterised by cribroperidinium spp. and sirmiodinium grossii. apteodinium spp., cassiculosphaeridium magna, paragonyaulacysta capillosa and p. borealis are locally common. the assemblage is recorded as moderately diverse and abundant. biostratigraphy. the p. scitulus ammonite zone has never been proven by ammonites in greenland but is included in the zonal scheme due to the general resemblance of the ammonite successions in greenland and england (fig. 14; birkelund et al. 1984). hence, we refer to it with caution, as indicated by the brackets. the top of abundant o. patulum occurs in ammonite faunal horizon m 25 (piasecki 1996). this succession therefore correlates with the (p. scitulus –) p. wheatleyensis chronozones. age. early volgian, late jurassic. organic matter. the interval is characterised by abundant amorphous kerogen together with a large proportion of terrestrial organic material from higher land plants, especially brown to black woody material. the presence of dinoflagellate cysts indicates a marine environment, and the dominance of aom is suggestive of oxygen-deficient bottom conditions. 5.5 pectinatites huddlestoni – pectinatites pectinatus chronozones undiff. (119.70–97.00 m) fossils. the recognition of the interval is based on its dinoflagellate cyst record. trichodinium piaseckii and senoniasphaera clavellii have the highest occurrence in 110.15 m, and the last occurrence of o. patulum is at 99.72 m near the top of the interval. the assemblage is characterised by oligosphaeridium patulum, paragonyaulacysta capillosa, p. borealis and sirmiodinium grossii. the assemblage is considered moderately diverse and abundant. within this interval, pterospermella spp. acritarchs have their first appearance and become common and abundant from this interval and upwards until their highest occurrence at 21.38 m. biostratigraphy. the last occurrences of trichodinium piaseckii and senoniasphaera clavellii are recorded between ammonite fauna horizons m 25 and m 29, p. wheatleyensis and p. pectinatus chronozones in milne land (fig. 14; piasecki 1996). this is in accordance with the type occurrences of the zonal index species in the north sea region, uk (bailey et al. 1997). the last consistent occurrence of o. patulum in milne land is a few metres above fauna m 25, p. wheatleyensis zone. age. early volgian, late jurassic. organic matter. abundant amorphous kerogen together with terrestrial organic material from higher land plants, especially sporomorphs and brown to black woody material, characterises the interval. the black woody material becomes more lath-shaped upwards in the succession. the presence of dinoflagellate cysts indicates a marine environment, and the dominance of aom is suggestive of oxygen-deficient bottom conditions. 5.6 dorsoplanites primus chronozone (and pavlovia iatriensis and pavlovia rugosa chronozones; 97.00–79.74 m) fossils. the recognition of this zone(s) is based on its content of dinoflagellate cysts and ammonites. the base is placed below the lowest occurrence of dorsoplanites ammonites (table 1) at the lithostratigraphic boundary between the bernbjerg and lindemans bugt formations (97 m). the interval marks the first appearance of ammonites after https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 22 of 39 geusbulletin.org an interval barren of ammonites between 199.42 m and 90.85 m. a crushed, medium-sized ammonite fragment at 90.85 m, showing subdued ribbing with straight primaries and intercalated secondaries beginning on mid-flank and a gently sloping umbilical wall, is assigned to dorsoplanites sp. (fig. 10b; table 1). a small fragment of a large specimen found at 89.30 m has straight, blunt distant ribs that bifurcate very high. the specimen is crushed but identified as the zonal index species d. primus by callomon & birkelund (1982; fig. 10c; table 1). in addition, mega onychites – arm hooks of belemnites – are encountered (fig. 15a; table 1). the palynomorph assemblage from one sample contains abundant dinoflagellate cysts: cassiculosphaeridium magna and leptodinium subtile and common cribroperidinium spp., paragonyaulacysta borealis, sirmiodinium grossii, tubotuberella apatela, as well as the acritarch pterospermella spp. the dinoflagellate cyst assemblage is relatively diverse and abundant, but chorate dinoflagellate cysts are essentially absent. biostratigraphy. dorsoplanites primus indicates the faunal horizon m 31 in milne land. the general appearance of the genus dorsoplanites in the east greenland ammonite succession is recorded in the d. primus zone and ranges up to m 42 in the upper middle volgian crendonites anguinus zone (fig. 14; callomon & birkelund 1982). the first occurrence of dorsoplanites encountered here agrees well with the presence of the d. primus zone. mega-onychites are generally considered indicative of upper jurassic deposits, kimmeridgian–volgian strata, in the arctic part of the boreal realm (hammer et al. 2013). the lack of age-diagnostic fossils in the upper part of this interval leaves room for the presence of the p. iatriensis and p. rugosa zones, which are thus indicated with caution. the dinoflagellate cysts cassiculasphaeridium spp. have their maximum abundance in this interval. biostrat (2018) reports a maximum occurrence of cassiculasphaeridium spp. in the middle of the subboreal p. pallasioides ammonite chronozone, which probably corresponds to the boreal d. primus – p. iatriensis ammonite chronozones (callomon & birkelund 1982). age. earliest middle volgian, late jurassic. organic matter. abundant amorphous kerogen together with terrestrial organic material from higher land plants, especially sporomorphs and brown to black woody material, characterises the interval. lath-shaped, black woody material is common. the presence of dinoflagellate cysts indicates a marine environment, and the dominance of aom is suggestive of oxygen-deficient bottom conditions. 5.7 pavlovia communis chronozone (79.74–74.20 m) fossils. the recognition of the zone is based on its content of dinoflagellate cysts and ammonites. the base is placed at the lowest appearance of muderongia simplex. one ammonite is recorded in the interval (75.43 m; table 1). the visible section of the ammonite shows coarse, biplicate rursiradiate ribbing from a large form, resembling various species within the genus pavlovia. it appears closest to p. variocostata callomon & birkelund, characterised by a modification to subdued, irregular widely spaced and extremely coarse biplicate ribs in adult stage. considering the restricted visible part, the present specimen is cautiously referred to as p. cf. variocostata (fig. 10d). the dinoflagellate cyst assemblage, based on one sample (79.74 m), contains apteodinium spp., cribroperidinium spp., muderongia simplex, sirmiodinium grossii, paragonyaulacysta borealis and p. capillosa. the diversity of the assemblage is relatively high, whereas the abundance is low. chorate cysts are almost absent. biostratigraphy. in north-west europe, the lowest appearance of common muderongia simplex spp. is well established in the p. rotunda (ammonite) zone (e.g. riding & thomas 1992; riding et al. 2000). the p. rotunda zone is correlated with the p. communis zone (m 34 to m 35) in east greenland (fig. 14; birkelund et al. 1984). the first appearance of muderongia simplex is thus here considered to indicate the p. communis zone. this agrees well with the presence of the ammonite p. cf. variocostata, recorded c. 4 m higher in the core. the ammonite indicates the m 35 faunal horizon, which lies in the upper part of the p. communis zone (fig. 14; callomon & birkelund 1982). the appearance of muderongia simplex in milne land occurs somewhat higher, in the faunal horizon m 46, in the c. anguinus ammonite zone (piasecki 1996) corresponding to its highest common occurrence in north-west europe (riding et al. 2000). age. middle volgian, late jurassic. organic matter. abundant amorphous kerogen together with terrestrial organic material from higher land plants, fig. 15 onychites in rødryggen-1 core. a: mega-onychites, level 89.26 m, mguh 34183. b: micro-onychites, level 62.43 m, mguh 34188. a b 1 cm https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 23 of 39 geusbulletin.org mainly spores and pollen and brown to black woody material, characterises the interval. lath-shaped, black woody material is common. the presence of dinoflagellate cysts indicates a marine environment, and the dominance of aom is suggestive of oxygen-deficient bottom conditions. 5.8 dorsoplanites liostracus chronozone (74.20–70.15 m) fossils. the base of this zone is placed at the occurrence of an ammonite specimen. no samples were analysed for dinoflagellate cysts within this interval. two ammonite specimens were found at closely spaced levels (table 1). the lower ammonite, a large form, shows dense, subdued ribbing, with blunt primaries that divide almost in a fasciculate manner into 3–4 secondaries and intercalatories. the subdued ribbing somewhat resembles, but does not exactly match, that of dorsoplanites liostracus callomon & birkelund. the present specimen is thus tentatively referred to dorsoplanites aff. liostracus (fig. 10e). the higher specimen is a crushed fragment with slightly flexuous, relatively strong, blunt primary ribs and weak intercalating secondaries (fig. 10f). this ribbing style resembles that of the much larger pavlovia corona callomon & birkelund (1982, pl. 3, fig. 1). the present specimen is thus referred to p. cf. corona. biostratigraphy. both p. aff. liostracus and p. cf. corona indicate the faunal horizon m37 in the upper part of the d. liostracus zone (fig. 14; callomon & birkelund 1982). age. middle middle volgian, late jurassic. organic matter. no palynological samples. 5.9 dorsoplanites gracilis chronozone (70.15–59.85 m) fossils. the base of this interval is placed at the lowest occurrence of the dinoflagellate cyst lagenorhytis delicatula. the interval contains both ammonites and dinoflagellate cysts. a few centimetres above the base of the zone, an ammonite fragment from an apparently evolute form shows ornamentation with relatively distant ribs (fig. 10g). primaries are blunt, gently prorsiradiate and forward concave. secondaries mostly intercalate and develop at mid-flank. considering the presence of p. cf. corona in the underlying zone and e. cf. pseudapertum in the overlying zone, the closest resembling ammonite forms are likely to be found in the dorsoplanites gracilis group in the faunal horizons m 37 to m 42 and are assigned to the d. liostracus – e. pseudapertum zones, middle volgian east greenland (spath 1936; callomon & birkelund 1982). the ribbing pattern of the present specimen resembles the ‘indistinct’ pattern of dorsoplanites jamesoni spath (1936, pl. 29, fig. 3). additional fossils include a crushed indeterminate ammonite, a poorly preserved recrystallised belemnite rostrum and small onychites (fig. 15b; table 1). the dinoflagellate cyst assemblage is relatively diverse and moderately abundant. it is dominated by lagenorhytis delicatula, the acritarch pterospermella spp. and cribroperidinium spp. chorate cysts are absent. biostratigraphy. the exact assignment of the species dorsoplanites jamesoni to a faunal horizon remains, but it probably belongs to the upper part of the d. gracilis zone and possibly the m 40 faunal horizon (fig. 14; birkelund et al. 1984). the first appearance of lagenorhytis delicatula is just below the occurrence of the ammonite dorsoplanites jamesoni at 70.06 m. l. delicatula is generally assigned a lower cretaceous range (costa & davey 1992) and is also present here in ryazanian strata higher in the core. hence, its first appearance in the middle volgian d. gracilis zone and consistent presence into the overlying ryazanian are surprising, and its first occurrence in the middle volgian potentially represents an excellent local marker event. age. middle middle volgian, late jurassic. organic matter. the interval is characterised by abundant amorphous kerogen, together with limited terrestrial organic material, mainly as spores and pollen, black woody material and carbonised, rounded grains. the presence of dinoflagellate cysts indicates a marine environment, and the dominance of aom is suggestive of oxygen-deficient bottom conditions. 5.10 epipallasiceras pseudapertum chronozone (59.85–54.53 m) fossils. the base of this zone is placed at the uppermost occurrence of leptodinium subtile. the interval contains both ammonites and dinoflagellate cysts. a small ammonite with well-preserved, relatively distant ribbing has strong, straight, prorsiradiate primaries that divide rather high up on the flank into two secondaries (fig. 10h). it resembles epipallasiceras pseudapertum spath (1936; pl. 9, fig. 4) and is cautiously referred to e. cf. pseudapertum. a level less than one metre higher contains another, but poorly preserved and indeterminate, ammonite (table 1). the dinoflagellate cyst assemblage, based on three samples, is abundant and diverse. it is dominated by cribroperidinium spp. with common lagenorhytis delicatula and apteodinium daveyi in the lower part. rare, small and poorly preserved gochteodinia villosa villosa and gochteodinia spp. occur at 56.88 m but are not https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 24 of 39 geusbulletin.org recorded in any samples above until 37.71 m, in the subcraspedites primitivus zone (upper volgian). chorate and cavate cysts are rare in the lower sample and slightly more common in the upper sample. biostratigraphy. the ammonite indicates the m-42 faunal horizon and is the index of the middle volgian e. pseudapertum zone (fig. 14). this is in good biostratigraphic accordance with the record of the uppermost occurrence of the dinoflagellate cyst l. subtile, which is a well-known marker from north-west europe, where it has its highest occurrence in the p. albani zone (riding & thomas 1992). the upper p. albani zone is correlated with the e. pseudapertum zone in the east greenland ammonite zonation (birkelund et al. 1984). the first appearances of perisseiasphaeridium insolitum and apteodinium daveyi at 59.85 m are slightly earlier than the recorded first occurrence of both species in the galbanites okusensis – g. kerberus ammonite zones in england (davey 1982). the dinoflagellate cyst assemblage (59.85–55.88 m) is dominated by cribroperidinium spp. but becomes poorer upwards. age. middle middle volgian, late jurassic. organic matter. this interval records a significant change to a low content of organic matter in general, especially of terrestrial woody material; this shift appears in the lower e. pseudapertum chronozone and continues upwards to the subcraspedites primitivus chronozone in the upper volgian. this is combined with low abundance and diversity of the dinoflagellate cysts assemblage. a fully marine environment is indicated by the fossil record. 5.11 laugeites groenlandicus chronozone (54.53–50.14 m) fossils. the zone is recognised by its ammonite assemblage, whereas the dinoflagellate cyst assemblage is impoverished and poorly preserved. acritarchs are present but have little stratigraphic value. the base of the zone is placed at the lowest occurrence of a possible laugeites ammonite (54.53 m), a few centimetres below a fragment of an apparently very large laugeites (table 1). the latter is characterised by straight, strong and distant primary ribs, which begin sharp but appear to become blunt and flatten on or towards the (mid?-)flank (fig. 10i). it resembles laugeites biplicatus mesezhnikov figured by repin et al. (2006), although that specimen is somewhat different from the holotype established by mesezhnikov (in zakharov & mesezhnikov 1974). considering its fragmented nature, the specimen from the core is cautiously referred to l. cf. biplicatus. further levels with ammonites include l. cf. intermedius donovan, another l. cf. biplicatus and l. cf. planus mesezhnikov (in zakharov & mesezhnikov 1974; table 1). the former closely resembles the specimen figured by donovan (1964, pl. 1, fig. 5) from laugeites ravine, kuhn ø (fig. 10j). the second is a fragment with distant, low, blunt ribs, which resembles the ornamentation on the mid-flank in the outer whorls of laugeites biplicatus as seen in, for example, repin et al. (2006, pl. 47, fig. 2) and rogov (2010, pl. 4, fig. 6). the latter is a complete, small juvenile (?) specimen with relatively high whorl sides and narrow umbilicus, very dense and delicate ribbing with more than 31 primaries per whorl that divide or bifurcate almost immediately or low on the flank into dense and fine secondaries (fig. 10k). ribbing is slightly concave. by comparison with material figured by donovan (1964) and surlyk et al. (1973), the specimen is referred to the genus laugeites, which is characterised as being smaller, more compressed and more delicately ribbed than its predecessor dorsoplanites (see discussion in description of dorsoplanites intermissus in callomon & birkelund 1982, appendix, p. 368). l. parvus donovan and other greenland laugeites closely resemble the specimen but have more forward-leaning ribs, whereas the ribbing on the present specimen is almost radiate. the closest resemblance of our specimen is with laugeites planus mesezhnikov (in zakharov & mesezhnikov 1974) from subarctic urals, russia. in addition, the interval contains indeterminate, crushed ammonite fragments and aptychi (table 1). the dinoflagellate cyst assemblage (samples from 54.31–50.53 m) is poor in the lower part of the zone with common apteodinium, cribroperidinium and lagenorhytis. the acritarchs cymatiosphaera, veryhachium, pterospermella, leiosphaeridia and the prasinophyte algae tasmanites occur scattered amongst the dinoflagellate cysts and become common in the upper part of the zone where dinoflagellate cysts disappear. biostratigraphy. in north-east greenland, the genus laugeites is restricted to the m 47 faunal horizon in the laugeites groenlandicus zone (fig. 14; spath 1936; donovan 1964; surlyk 1978; callomon & birkelund 1982). the present records of l. cf. planus and l. cf. biplicatus are new for greenland, where l. groenlandicus (spath 1936), l. parvus donovan (1964) and l. intermedius donovan (1964) had been recorded previously. l. biplicatus indicates the epivirgatites nikitini zone in russia, which is correlative to the upper part of the laugeites groenlandicus zone and the overlying epilaugeites vogulicus zone sensu surlyk (1978), i.e. m 47 and higher levels (rogov 2010, 2020). in the rødryggen-1 core, l. cf. biplicatus occurs below l. cf. planus, which indicates the laugeites groenlandicus zone and m 47 faunal horizon; this delimits the occurrence of l. cf. biplicatus in the core to the https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 25 of 39 geusbulletin.org l. groenlandicus zone. note that kelly et al. (2015) and rogov (2020), based on observations of the ammonite faunal succession in perisphinctes ravine, eastern kuhn ø (fig. 2), added another faunal horizon, the l. lambecki horizon, above the m 47 laugeites groenlandicus horizon, to the l. groenlandicus zone. however, callomon & birkelund (1982) argued that the laugeites fauna recorded in the northern wollaston forland and kuhn ø areas by donovan (1964) and surlyk (1978) represents the higher part of the l. groenlandicus zone. it seems more justified to introduce a l. parvus horizon, based on donovan’s and surlyk’s documentation of the fauna, above the m 47 horizon, rather than an l. lambecki horizon and zone. the absence of crendonites anguinus zone fossils is not necessarily attributed to a hiatus, given that there are no sedimentological or depositional signs of erosion. the zone is thus considered not proven in the core, either due to the low sample density or due to low sedimentation rates. age. late middle volgian. organic matter. the interval contains very little organic matter with few or no dinoflagellate cysts. despite the dominance of acritarchs and terrestrial organic material, the presence of ammonites testifies to continued fully marine conditions. 5.12 praechetaites exoticus chronozone (50.14–45.37 m) fossils. the zone is recognised by its ammonite assemblage, whereas the dinoflagellate cyst assemblage is poor and poorly preserved. acritarchs are present but have little stratigraphic value. the interval is relatively rich in macrofossils containing several ammonites, various bivalves, including buchia, and ammonite aptychi. the base is placed at 50.14 m, at the occurrence of a crushed ammonite fragment with relatively distant, low, sharp, primary ribs that divide into three closely spaced low, sharp, secondary ribs (table 1). the ornamentation is close to that in inner–middle whorls of praechetaites exoticus (shulgina) as illustrated in, for example, shulgina (1967, pl. 1, fig. 1b). an ammonite occurring c. 1.5 m higher is a fragment, probably from mid-flank, with very faint ribbing of distant primaries that are very low, almost smooth, developing into faint secondaries and intercalatories (fig. 10l) similar to ornamentation in the mid-flank of praechetaites exoticus (shulgina; e.g. shulgina 1967, pl. 4, fig. 1; rogov 2010, pl. 6, fig. 5). a few centimetres higher in the core, another ammonite fragment has low, wide, relatively sharp crested ribs that appear to develop into faint sheaves of very fine lirae-like secondaries (fig. 10m), resembling the ornamentation observed in praechetaites exoticus (shulgina; e.g. shulgina 1967, pl. 4, fig. 1). considering their preservation, the fragments are with some caution referred to praechetaites cf. exoticus (shulgina). two additional ammonites are indeterminate fragments, one of them possibly also belonging to the genus praechetaites (fig. 10n; table 1). the dinoflagellate cyst assemblage varies from barren to poor. reworked specimens become more common, with specimens of wanaea and nannoceratopsis. acritarchs are present in all samples, for example, fromea, pterospermella and leiosphaeridia. biostratigraphy. p. exoticus is the index species of the northern siberian p. exoticus zone. its presence in greenland closely above the l. groenlandicus zone does not leave much room for the epilaugeites vogulicus zone that supposedly overlies the laugeites groenlandicus zone (fig. 14; surlyk 1978). it could thus be that the p. exoticus and e. vogulicus zones may be time-equivalent correlatives. however, since representatives of the two zones were so far not found together, it remains uncertain whether p. exoticus occurs in the e. vogulicus zone in greenland and thus would link the russian and greenland zones, or whether the p. exoticus zone may represent a level in between the l. greenlandicus zone and the e. vogulicus zone. rogov (2010, 2020) correlated the p. exoticus zone to both the e. vogulicus zone and the p. tenuicostatus zone or beds in greenland. rogov & zakharov (2011) noted that the specimens of e. vogulicus from greenland (surlyk et al. 1973; surlyk 1978) differ from ‘true’ e. vogulicus in siberia (ilovaisky 1917; mikhailov 1966; zakharov & mesezhnikov 1974). accordingly, the greenland records were referred to a new species, e. surlyki rogov (2020), which thus also becomes index for the e. surlyki zone in greenland. age. latest middle volgian. it should be noted, however, that there is disagreement amongst russian ammonite stratigraphers about whether the p. exoticus zone should be attributed to the uppermost middle or the lowermost upper volgian (meledina et al. 2010; rogov & zakharov 2011). the unit was originally introduced as a subzone in the lowest part of the craspedites okensis zone (upper volgian), whereas the ammonite succession in the jurassic–cretaceous boundary key section at nordvik suggests that the p. exoticus zone should be referred to the uppermost middle volgian (e.g. zakharov & rogov 2006, 2008; rogov & zakharov 2009, 2011; rogov 2020). organic matter. the interval has a very low organic content. marine palynomorphs are rare or absent, although acritarchs are present. they contain a limited component of terrestrial organic material, mostly sporomorphs and unstructured organic material. woody https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 26 of 39 geusbulletin.org material is mostly carbonised and degraded to small angular fragments. although the palynomorphs provide a predominantly terrestrial signal, ammonites are present indicating a fully marine environment. 5.13 subcraspedites primitivus chronozone (45.37–37.49 m) and lowermost gochteodinia villosa villosa zone (neg cr 1; 37.71–37.49 m) fossils. the interval is recognised from its ammonite and dinoflagellate cyst content. the base of the s. primitivus zone is located at the occurrence of an ammonite fragment showing strong, distant, bullate primaries that divide into somewhat weak secondaries (fig. 10p; table 1). these sets of secondaries are intercalated with weak ribs that start at a level equal to the furcation level of the primaries, resulting in an appearance of distant primaries and dense secondaries and intercalatories. the bullate primaries are forward-leaning, whereas the rear secondary in the rib sets is relatively rectiradiate on the mid-flank before leaning forward towards the ventral shoulder. this gives a sinuous appearance of the ornamentation. the closest resemblance is found outside greenland in the taxon subcraspedites (swinnertonia) subundulatus, known from england (casey 1973). it has similar distant primaries and dense secondaries with a sinuous appearance. the limited material only allows cautious reference to s. (s.) cf. subundulatus. another fragment of a small ammonite, a couple of metres higher in the core, shows dense, fine, delicate parallel ribbing (fig. 10q; table 1). its reference to subcraspedites (swinnertonia) sp. juv. rests both on the other specimen described from the zone and on comparison with the similarly finely ribbed juvenile subcraspedites (swinnertonia) sp. juv. from england (casey 1973, pl. 4). in addition, the interval contained another ammonite fragment, poorly preserved and not sampled, and a buchia bivalve. dinoflagellate cysts are rare or absent. relatively common taxa are apteodinium daveyi, cribroperidinium spp., cassiculosphaeridium magnum and sirmiodnium grossii. other species are present but very rare in one or two samples, whereas acritarchs, especially pterospermella spp., are common to abundant in a thin interval (40.62–38.28 m). stratigraphically significant species are generally missing, but the appearance of gochteodinia villosa villosa in the s. primitivus chronozone indicates the lower boundary of the g. villosa villosa zone (neg cr 1; nøhr-hansen et al. 2020) and is followed by consistent g. villosa villosa in higher strata. this event is stratigraphically slightly higher (one ammonite zone) than the reported first appearances in the subboreal region (paracraspedites oppressus zone; e.g. woollam & riding 1983), but, in contrast, the odd occurrence of rare g. villosa villosa mentioned above in the e. pseudapertum chronozone (59.85–54.53 m) is much lower. from the russian platform, riding et al. (1999) report the first appearance of rare g. villosa in the kachpurites fulgens ammonite zone, which they correlate with the subcraspedites preplicomphalus ammonite zone of north-west europe (rogov 2020). presumed reworked dinoflagellate cysts are present, for example, ambonosphaera staffinensis, gonyaulacysta jurassica, endoscrinium galeritum, nannoceratopsis sp., pareodinia halosa and taeniaesporites iunctispina, mostly derived from oxfordian–kimmeridgian strata. biostratigraphy. subcraspedites (swinnertonia) subundulatus occurs in the upper volgian subcraspedites (swinnertonia) primitivus zone in england (casey 1973, pl. 4, fig. 1). that zone is now adopted in the east greenland zonation (fig. 14). the upper volgian is poorly represented by ammonites in greenland, and the upper jurassic key section in milne land has a hiatus between the middle volgian epilaugeites surlyki zone and the lower valanginian (callomon & birkelund 1982). a more complete volgian succession appears to occur in the wollaston forland area. in addition to the s. primitivus zone recorded here in the rødryggen-1 core, rogov (2010, 2020) recognised supracraspedites sowerbyi in the faunal succession, presumably in the casp collection from perisphinctes ravine, eastern kuhn ø (fig. 2; kelly et al. 2015). the specimens were assigned to ‘beds with s. sowerbyi’. the species as such indicates the s. preplicompalus zone in england, where it overlies the s. primitivus zone (casey 1973). rogov (2010, 2020) also identified beds with chetaites chetae overlying subcraspedites sowerbyi. the upper volgian may thus possibly be subdivided into s. primitivus – s. preplicomphalus – c. chetae zones. the gochteodinia villosa villosa zone (nøhr-hansen et al. 2020) is revised. the lower boundary is now defined by the first occurrence of common presence g. villosa villosa. it was previously defined by the first occurrence of the taxon, but random records of g. villosa villosa stratigraphically lower than the g. villosa villosa zone in the east greenland successions makes the former definition inaccurate. age. late volgian. organic matter. the organic content is low, and terrestrial woody material is rare. the diversity and abundance of marine plankton is low in this zone and becomes significantly reduced up through the zone; the lowest content is recorded in the uppermost sample (38.28 m). 5.14 praetollia maynci chronozone and lower gochteodinia villosa villosa zone (neg cr 1; 37.49–27.73 m) fossils. the interval is recognised from its ammonite and dinoflagellate cyst content. the base of the zone is placed https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 27 of 39 geusbulletin.org at the occurrence of an ammonite fragment (37.49 m) with delicate, low, sharp, slightly sinuous ribs (fig. 10r; table 1). its ornamentation resembles praetollia maynci spath as illustrated in, for example, spath (1952) and surlyk (1978, pl. 5, fig. 1). a buchia bivalve was also recorded in this interval. four samples with abundant and diverse dinoflagellate cysts assemblages are placed in this ammonite chronozone. sirmiodinium grossi and cribroperidinium spp. dominate the assemblage together with common cassiculosphaeridium magnum, gochteodinia villosa villosa and paragonyaulacysta borealis. the acritarch pterospermella spp. is common. istmocystis distincta appears for the first time at 35.02 m, and canningia compta, scriniodinium pharo and palaecysta palmula have their first appearances higher in the zone. biostratigraphy. p. maynci is the index species of the p. maynci zone (fig. 14). the base of the ryazanian is the base of the cretaceous in a boreal sense. however, a formally chosen base of the cretaceous, in terms of the global boundary stratotype section and point (gssp) of the berriasian stage in the tethyan realm, is yet to be defined. it has recently been proposed to take the base of the calpionella alpina calpionellid zone as the boundary, which would most probably correlate to the russian taimyroceras taimyrensis and c. nodiger zones and the english s. lamplughi zone in the boreal upper volgian stage (gale et al. 2020; wimbledon et al. 2020). age. earliest ryazanian, early cretaceous. organic matter. the organic content is low, and terrestrial woody material is rare. in contrast, the dinoflagellate cyst assemblage is rich and diverse, in contrast to the poor assemblages in the underlying s. primitivus chronozone, and indicates fully marine conditions. 5.15 hectoroceras kochi chronozone and lower gochteodinia villosa villosa zone (neg cr 1; 27.73–25.82 m) fossils. the interval is recognised from its ammonite and dinoflagellate cyst content. an ammonite with ribs that radiate from a very narrow umbilicus indicates pronounced involuteness (fig. 10s; table 1). primaries divide into two secondaries and bend forward from the level of furcation. this ornamentation and the narrow umbilicus characterise the genus hectoroceras, and the specimen is referred to hectoroceras sp. regarding the dinoflagellate cysts (one sample), palaecysta palmula dominates the unit. the acritarch pterospermella spp. is common. the assemblage is relatively diverse and abundant. biostratigraphy. species of the ammonite genus hectoroceras are only recorded in the hectoroceras kochi zone (fig. 14; wright et al. 1996). the base of the zone is placed at the ammonite in the core at 27.73 m. the top of the zone is based at the last occurrence of rotosphaeropsis thule (g. villosa zone, top of dsk1 subzone; poulsen & riding 2003). palaecysta palmula appears in the top of the h. kochi zone of the north-west european subboreal ammonite zonation (riding & thomas 1992). in the rødryggen-1 core, p. palmula appears just above the ammonite. the p. maynci and h. kochi ammonite zones correlate with the lower g. villosa villosa zone (nøhr-hansen et al. 2020). the hectoroceras kochi zone is the highest ammonite zone identified and is correlated with the dinoflagellate cyst stratigraphy in this core. alsen (2006) identified ammonites and ammonite zones in the albrechts bugt member in sections near the rødryggen-1 core. due to solifluction in the outcrop sections, precise correlation with the core is not possible. correlation of the dinoflagellate cyst stratigraphy with local ammonite stratigraphy in the upper ryazanian to valanginian is therefore beyond the scope of this study. the ryazanian fauna is strictly boreal, whereas the valanginian fauna comprises tethyan, subboreal and boreal elements (alsen 2006). age. latest early ryazanian, early cretaceous. organic matter. the organic content is low. degraded organic material dominates, terrestrial material is represented by sporomorphs and woody material is very rare. the fossil record testifies to a fully marine environment. 5.16 upper gochteodinia villosa villosa zone (neg cr 1; 25.82–17.82 m) fossils. the interval is recognised based on its content of dinoflagellate cysts from the last occurrence of r. thule at 25.82 m to the first appearance of oligosphaeridium complex at 17.82 m. the last occurrences of paragonyaulacysta borealis and p. capillosa are in this interval. epiplosphaera spp., heterosphaeridium? spp., sirmiodinium grossii, palaecysta palmula and systematophora daveyi dominate this unit locally. the assemblage is relatively diverse and abundant in the lower part but becomes poorer in the upper part. this change is associated with a significant depositional shift from black shale to calcareous mudstone, probably reflecting a shift to more oxygen-rich bottom conditions. biostratigraphy. r. thule has its last occurrence in the h. kochi zone of the subboreal ammonite zonation in north-west europe (riding & thomas 1992; biostrat 2018). o. complex appears for the first time at the ryazanian–valanginian boundary in north-west europe (riding & thomas 1992) and in the peregrinoceras albidum zone, uppermost ryazanian, on store koldewey, northeast greenland (nøhr-hansen et al. 2020). https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 28 of 39 geusbulletin.org age. late ryazanian, early cretaceous. organic matter. the organic content is very low and totally dominated by black, rounded grains, probably reflecting strongly oxidising bottom conditions. the presence of marine plankton, however, is indicative of a fully marine environment. 5.17 lower oligosphaeridium complex zone (neg cr 2; 17.82–10.17 m) fossils. the interval is recognised based on its content of dinoflagellate cysts in the succession from the first appearance of oligosphaeridium complex (15.85 m) to the first occurrence of nelchinopsis kostromiensis. oligosphaeridium complex, oligosphaeridium spp. and epiplosphaera spp. dominate the assemblage. the assemblage is relatively diverse and abundant in the lower part but becomes poorer upwards. the interval is also represented by two samples analysed for calcareous nannofossils (at 16.6 and 13.82 m). the low diversity assemblage in the lower sample is dominated by watznaueria barnesiae. watznaueria fossacincta and crucibiscutum salebrosum are also common. in addition, watznaueria britannica, watznaueria ovata, retecapsa schizobrachiata, retecapsa crenulata, rotelapillus crenulatus, sollasites horticus, staurolithites stradneri, assipetra infracretacea, tegumentum stradneri, rhagodiscus asper, zeugrhabdotus fissus and nannoconus concavus are present. the upper sample has an assemblage dominated by crucibiscutum salebrosum and watznaueria barnesiae. also present are assipetra infracretacea, diazomatolithus lehmanii, retecapsa crenulate, rhagodiscus asper, watznaueria fossacincta, watznaueria britannica, watznaueria manivitiae, retecapsa angustiforata, zeugrhabdotus embergeri, cyclagelosphaera margerelii, staurolithites stradneri, stradnerlithus silvaradius, tegumentum stradneri, zeugrhabdotus erectus, tripinnilithus shetlandensis, zeugrhabdotus fissus and thoracosphaera spp. biostratigraphy. o. complex appears for the first time at the ryazanian–valanginian boundary in north-western europe (riding & thomas 1992), and in the top ryazanian of north-east greenland, peregrinoceras albidum zone (stefan piasecki unpublished data). the highest occurrence of gochteodinia villosa villosa and palaecysta palmula is in the lower valanginian (davey 1982; riding & thomas 1992). poulsen & riding (2003) correlate the last occurrence of g. villosa villosa with the top of the p. albidum zone in north-west europe, but this taxon is present into the lower valanginian in the rødryggen-1 core. the highest occurrence of palaecysta palmula is at the top of the paratollia zone, upper lower valanginian, in north-west europe (costa & davey 1992) and at the top of the ‘ekz6 zone’ in biostrat (2018; mid-lower valanginian). nelchinopsis kostromiensis is first recorded at 10.17 m, corresponding to the last occurrence of p. palmula (costa & davey 1992). however, in north-east greenland, n. kostromiensis occurs for the first time in the upper valanginian (nøhr-hansen et al. 2020). the calcareous nannofossil stratigraphy agrees fairly well with the palynostratigraphy. it indicates late ryazanian to early valanginian ages from high abundances of c. salesbrosum and watznaueria spp. co-occuring with s. horticus in the upper ryazanian to lower valanginian of the north sea area (jeremiah 2001) and offshore mid-norway (mutterlose & kessels 2000). r. crenulata has its first occurrence in the upper upper ryazanian (jakubowski 1987). the presence of nannoconus concavus and tripinnilithus shetlandensis in the absence of micrantholithus speetonensis indicates the lower valanginian nannofossil subzone bc3b (bown et al. 1998). pauly et al. (2012a) documented the first occurence (fo) of crucibiscutum spp., rhagodiscus asper and watznaueria spp. in the upper ryazanian of north-east greenland. age. early valanginian, early cretaceous. organic matter. the organic content is very low. the presence of marine plankton, albeit scarce, indicates marine conditions. the organic terrestrial material is heavily oxidised into carbonised, angular to rounded black grains. 5.18 upper oligosphaeridium complex zone (neg cr 2; 10.17–1.10 m) fossils. the interval is recognised based on its content of dinoflagellate cysts. the base is placed at the first occurrence of nelchinopsis kostromiensis (10.17 m), followed by the last occurrences of palaecysta palmula (9.00 m) and g. villosa villosa (6.59 m), and the highest record of lagenorhytis delicatula (1.10 m). the top of the zone is not recorded. the dinoflagellate cyst assemblage is very poor, and the presence or absence of several species may reflect random occurrences. however, oligosphaeridium complex dominates the assemblage, and cassiculosphaeridium magnum, circulodinium distinctum, epiplosphaera spp. and downiesphaeridium tribuliferum are common locally. the dinoflagellate cysts are fairly well preserved but constitute a very small fraction of the organic matter. the interval also contains one sample analysed for its calcareous nannofossil content. the nannofossil assemblage has a low diversity dominated by watznaueria barnesiae. also present are watznaueria fossacincta, calculites? sp.1, tranolithus gabalus, staurolithites stradneri, rhagodiscus asper, watznaueria britannica and thoracosphaera spp. c. salesbrosum is absent. biostratigraphy. the occurrence of nelchinopsis kostromiensis indicates a late valanginian age based on reports from the arctic (davies 1983; nøhr-hansen et al. 2020; ingrams et al. 2021) in contrast to north-west europe where it appears in lower valanginian strata (e.g. heilmann-clausen https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 29 of 39 geusbulletin.org & birkelund 1987; costa & davey 1992; biostrat 2018). the last occurrences of gochteodinia villosa villosa and p. palmula and the continuous presence of lagenorhytis delicatula in the uppermost core sample may suggest upper valanginian strata as recorded in north-west europe (davey 1982; riding & thomas 1992). the calcareous nannofossils t. gabalus, r. asper and watznaueria spp. range from the upper ryazanian in the north sea (jakubowski 1987; jeremiah 2001), north-east greenland (pauly et al. 2012a) and off mid-norway in the northern north atlantic (mutterlose & kessels 2000). calculites? sp.1 was described in bown et al. (1998) as ranging from the early hauterivian? to the early? barremian, albeit in tunisia and bulgaria. of note is the absence of c. salesbrosum. in north-east greenland, this species has its last appearance datum in the late hauterivian (pauly et al. 2012a) and a little later in the north sea (jakubowski 1987). however, the absence of this species is probably an artefact of preservation, since the age suggested by its apparent absence is too young compared with the palynostratigraphy of the interval, as well as the ammonite stratigraphy obtained in an outcrop study of the albrechts bugt member at the drill site (alsen 2006). age. late valanginian, early cretaceous. organic matter. the interval is characterised by a low organic content with little marine plankton and no fig. 16 the ranges of key dinoflagellate cyst taxa in the brorson halvø-1 core. the dinoflagellate species are arranged according to the succession of their lowest occurrences. question mark (?): uncertainty in taxonomic identification. r: reworked or redeposited. a range chart showing all taxa recorded in the core is available in supplementary data file 2. h-eb: hauterivian – early barremian. ler: latest early ryazanian. bl-nk: b. longicornutum zone – n. kostromiensis subzone. e.: early. m.: middle. 5 10 15 20 25 30 35 40 45 50 55 60 65 70 75 80 85 90 95 100 105 110 115 120 125 130 135 140 145 150 155 160 165 170 175 180 185 190 195 200 205 210 215 220 225 230 td stratumbjerg palnatokes bjerg lindemans bugt b er nb je rg rødryggen albrechts bugt storsletten boreal early cretaceous b or ea l l at e ju ra ss ic e. barremian h-eb ler m. volgian e. v ol gi an ki m m er id gi an h. kochi e. pseudapertum – d. liostracus p. wheatleyensis p. e le ga ns a . a ut is si od or en sis b. longicornutum zone, p. anaphrissum subzone bl-nk g. villosa villosa zone pe ri ss ei as ph ae ri di um p an no su m sc ri ni od in iu m in ri ti bi le o lig os ph ae ri di um p at ul um c ri br op er id in iu m c om pl ex um en do sc ri ni um lu ri du m le pt od in iu m su bt ile tr ic ho di ni um p ia se ck ii se no ni as ph ae ra c la ve lli g oc ht eo di ni a m ut ab ili s w al lo di ni um k ru tz sc hi i la ge no rh yt is de lic at ul a g oc ht eo di ni a vi llo sa v ill os a ps eu do ce ra tiu m p el lif er um m ud er on gi a st au ro ta m ud er on gi a te tr ac an th a n el ch in op sis k os tr om ie ns is eg m on to di ni um to ry nu m ps eu do ce ra tiu m a na ph ri ss um r ? r ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? b ar re n 3.37 6.16 7.03 9.18 10.30 13.05 15.87 20.68 22.22 27.85 33.41 34.43 35.40 36.36 37.35 37.52 38.30 39.29 40.26 41.28 42.09 45.29 46.05 49.70 56.19 65.23 76.28 85.77 101.00 111.25 120.80 136.40 149.83 165.85m 180.06m 195.73m 205.20m 214.98m 219.73m 223.73m 225.88m 3.37 top pseudoceratium anaphrissum 6.16 top nelchinopsis kostromiensis 9.18 base muderongia staurota base muderongia tetracantha 37.35 base pseudoceratium pelliferum base egmontodinium torynum base gochteodinia villosa base systematophora palmula 38.30 top leptodinium subtile 42.09 base lagenorhytis delicatula 46.05 top senoniasphaera clavelli 49.70 top abundant oligosphaeridium patulum 56.19 base rhynchodiniopsis martonense 65.23 top gochteodinia mutabilis 76.28 base gochteodinia mutabilis base wallodinium krutzschii 85.77 top endoscrinium luridum 111.25 base senoniasphaera clavelli base trichodinium piaseckii top cribroperidinium complexum 195.73 base oligosphaeridium patulum top scriniodinium irregulare top perisseiasphaeridium pannosum pe ri od / ep oc h m em be r fo rm at io n d ep th (m ) lithostratigraphy a ge zo ne zo ne chronostratigraphy d in o� ag el la te c ys t z on es (n øh rh an se n et a l. 20 20 ) a m m on it e ch ro no zo ne s (c al lo m on & b ir ke lu nd 19 82 ) sa m pl e de pt h is b a se of d ep th ra ng e (m ) samples events dino�agellate cysts stratigraphic range a . e ud ox us 149.83 https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 30 of 39 geusbulletin.org recognisable terrestrial organic material. marine conditions are indicated by the macrofauna and calcareous flora. 6. biostratigraphy of the brorson halvø-1 core the biostratigraphic subdivision of the brorson halvø-1 core is described from td at 225.7 m upwards. the location of bulk-rock samples is seen in the charts, illustrating the distributions and ranges of dinoflagellate cyst and calcareous nannofossil taxa (figs 16, 17). selected dinoflagellate cysts are figured in fig. 13. 6.1 aulacostephanus eudoxus chronozone (225.7 m (td) – 195.73 m) fossils. the recognition of the zone is based on the dinoflagellate cyst record. the base of the zone is placed at the base of the core (at td 225.7 m). the dinoflagellate cyst assemblage is generally poor, and dinoflagellate cysts are not common; gonyaulacysta jurassica, g. dualis, sirmiodinium grossii and scriniodinium irregulare are locally common. perisseiasphaeridium pannosum is rare to common, but not as abundant as normally recorded in east greenland. biostratigraphy. p. pannosum occurs in the lowermost analysed sample at 225.88 m and is common at the top, at 195.73 m, where it coincides with the appearance of common to abundant oligosphaeridium patulum. this indicates an interval corresponding to the ammonite horizons m 20 – m 22, a. eudoxus zone (piasecki 1996). no other recorded dinoflagellate cysts indicate correlation to zones below the a. eudoxus chronozone, and the succession correlates, therefore, with the a. eudoxus chronozone. the dinoflagellate cysts atopodinium haromense, dingodinium minutum, nannoceratopsis pelucida, paragonyaulacysta capillosa and taeniophora iunctispina occur scattered throughout this zone. age. kimmeridgian, late jurassic. organic matter. amorphous kerogene dominates together with terrestrial, organic material from higher land plants, especially sporomorphs and black woody material. the woody material is physically degraded to rounded and angular grains in the lower levels of the succession; the grains become lath-shaped and larger upwards. the presence of dinoflagellate cysts indicates a marine environment, and the dominance of aom is suggestive of oxygen-deficient bottom conditions. remarks. the interpreted late kimmeridgian age at the base of the well agrees with the find of an ammonite in a measured outcrop section, c. 200 m below the albrechts bugt member, palnatokes bjerg formation (m. bjerager, pers. comm. 2022). the ammonite is identified as amoeboceras subkitchini from the upper lower kimmeridgian rasenia cymodoce zone (figs 14, 18). fig. 17 the distribution of calcareous nannofossils in the brorson halvø-1 core. 5 10 15 20 25 30 35 40 not analysed barren bi sc ut um c on st an s c ru ci bi sc ut um sa le br os um d isc or ha bd us ig no tu s rh ag od is cu s a sp er tr iq ue et ro rh ab du lu s s he tla nd en sis w at zn au er ia b ar ne si ae et hm or ha bd us h au te ri vi an us h em ip od or ha bd us g or ka e re te ca ps a su ri re lla ze ug rh ab do tu s d ip lo gr am m us ze ug rh ab do tu s e re ct us st au ro lit hi te s c ru x w at zn au er ia sp p. li th ra ph id ite s c ar ni ol en sis so lla si te s h or tic us ei � el lit hu s s tr ia tu s st ra dn er lit hu s s ilv ar ad iu s cy cl ag el os ph ae ra m ar ge re lii tr an ol ith us g ab al us w at zn au er ia b ri ta nn ic a d ia zo m at ol ith us le hm an ii m ic ra nt ho lit hu s o bt us us te gu m en tu m o ct ifo rm is bu kr yl ith us a m bi gu us pe ri ss oc yc lu s p le th ot re tu s sp ee to ni a co lli ga ta ze ug rh ab do tu s s cu tu la pe rc iv al ia fe ne st ra ta pe ri ss oc yc lu s t ay lo ri ae ro te la pi llu s l a� tt ei ze ug rh ab do tu s t ri ve ct is 6.16m 7.03m 8.26m 9.18m 12.07m 14.89m 16.88m 17.65m 18.63m 22.22m 26.11m 30.52m 31.46m 32.44m 33.41m 34.43m 35.40m 37.35m 40.26m hauterivian d ep th (m ) lithostratigraphy chronostratigraphy samples nannofossil bc zonation (bown et al. 1998) b ar re n sa m pl e de pt h is b a se of d ep th ra ng e (m ) nannofossils semi-quantitative (default abundance scheme) pe ri od / ep oc h m em be r fo rm at io n a ge su bz on e zo ne stratumbjerg palnatokes bjerg lindemans bugt rødryggen albrechts bugt storsletten valanginian late hauterivian early hauterivian late valanginian bc8 bc6/?bc7 bc5 bc8c bc8a/bc8b upper bc5 lower bc5 https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 31 of 39 geusbulletin.org 6.2 aulacostephanus autissiodorensis chronozone (195.73–149.83 m) fossils. the recognition of the zone is based on its dinoflagellate cyst record. the base of the zone is placed at the first appearance of common to abundant oligosphaeridium patulum (195.73 m) and its top at the first appearance of trichodinium piaseckii coincident with the last occurrence of cribroperidinium complexum (149.83 m). the dinoflagellate cyst assemblage is generally poor but is characterised by common to abundant oligosphaeridium patulum and paragonyaulacysta capillosa. atopodinium spp., circulodinium spp., pareodinia spp. and sirmiodinium grossii are common locally. the assemblage is mostly of low diversity and low abundance, probably due to the abundance of organic matter in the samples. biostratigraphy. common to abundant o. patulum appears between ammonite horizons m 22 – m 23 in milne land near the boundary between the a. eudoxus and a. autissiodorensis zones (piasecki 1996; alsen & piasecki 2018). t. piaseckii appears below ammonite fauna horizon m 25, p. wheatleyensis zone (piasecki 1996). the succession is, therefore, correlated with the a. autissiodorensis and p. elegans zones. age. kimmeridgian, late jurassic. organic matter. the interval contains amorphous kerogen together with abundant terrestrial organic material from higher land plants, especially sporomorphs and black, lathshaped woody material. the presence of dinoflagellate cysts indicates a marine environment, and the dominance of aom is suggestive of oxygen-deficient bottom conditions. 6.3 pectinatites elegans chronozone (149.83–65.23 m) fossils. the zone is based on the dinoflagellate cyst record. the base of the zone is recognised at the first occurrence of trichodinium piaseckii coincident with the last occurrence of cribroperidinium complexum. the dinoflagellate cyst assemblage is dominated by cribroperidinium spp., circulodinium spp., oligosphaeridium patulum, pareodinia sp., paragonyaulacysta borealis and p. capillosa in the lower part of the zone. cassiculosphaeridium magnum, gochteodinia mutabilis, oligosphaeridium patulum, paragonyaulacysta capillosa, sirmiodinium grossii, tenua hystrix, tubotuberella spp. and the acritarch wallodinium krutschi are more common in the upper part of the zone. the assemblage is relatively diverse and abundant. biostratigraphy. the first occurrence of t. piaseckii is coincident with the last occurrence of c. complexum in this core and in the rødryggen-1 core, where it correlates with the lower boundary of the p. elegans zone. the first occurrence of t. piaseckii is followed successively by the first occurrences of senoniasphaera clavellii and gochteodinia mutabilis and the last occurrence of common o. patulum. all these events are located below ammonite horizon m 25, wheatleyensis zone, in milne land (piasecki 1996) suggesting correlation with the p. elegans zone. these events coincide in the relatively condensed succession in milne land, whereas they occur in succession in both rødryggen-1 and brorson halvø-1 cores. age. early volgian, late jurassic. organic matter. abundant amorphous kerogen dominates together with abundant terrestrial organic material from higher land plants, especially sporomorphs and brown and black woody material. the presence of dinoflagellate cysts indicates a marine environment, and the dominance of aom is suggestive of oxygen-deficient bottom conditions. 6.4 pectinatites wheatleyensis chronozone (65.23–45.5 m) fossils. the zone is based on the dinoflagellate cyst record. the base of this zone is placed at the last occurrence of common gochteodinia mutabilis. the top of the zone is bounded by an unconformity marking a significant hiatus. fig. 18 amoeboceras subkitchini (mguh 34201 from geus 469823) ammonite from upper lower kimmeridgian rasenia cymodoce zone. collected stratigraphically c. 200 m below the albrecht bugt member (palnatokes bjerg fm) in an outcrop section measured along a small ravine from the brorson halvø-1 drill site and towards the west. the specimen is housed in the palaeontology type collection at the natural history museum of denmark and labelled with an mguh number – museum geologica universitas hafniensis. 1 cm https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 32 of 39 geusbulletin.org fig. 19 the correlation between the rødryggen-1 and brorson halvø-1 boreholes; the datum is the base of the palnatokes bjerg formation. note the complete stratigraphy in the rødryggen-1 well compared to the significant hiatuses within the volgian–ryazanian in the brorson halvø-1 core. the zones that are present in the rødryggen-1 well and absent in the brorson halvø-1 well are here shown schematically to onlap the unconformity. although it could be argued that these zones were also represented in the brorson halvø-1 area and subsequently removed by erosion of the crest during tilting of the fault block, the lack of redeposited palynomorphs of these ages in the rødryggen-1 core favours an onlap model. ab: albrechts bugt mb. lb: lindemans bugt formation. pb: palnatokes bjerg fm. rr: rødryggen member. str: stratumbjerg formation. see fig. 4 for legend and grain size. p. an ap hr isu m n . k os tr om ie ns is bc 8 bc 6 (bc 7) h . k oc hi p. m ay nc i s. p rim iti vu s p. ex ot ic us l. g r p. co m m un is d . p rim us p. ia tr ie ns is p. ru go sa bc 5 an d g . v ill os a vi llo sa d . l io st ra cu s d . g ra cil is e . p se ud ap er tu m p. w hea tle ye nsis p. ele ga ns a . a ut iss io do re ns is a . e ud ox us a. eudoxus a. autiss iodorensis p. e leg ans p. whe at le ye ns is p. h ud dl es to ni p. pe ct in at us d . p rim us p. ia tr ie ns is p . r ug os a p. co m m un is l. g ro en la nd ic us p. ex ot ic us e. p se ud ap er tu m d . g ra ci lis s. p rim iti vu s a nd lo w er m os t g . v ill os a vi llo sa p. m ay nc i a nd lo w er g . v ill os a vi llo sa up pe r g . v ill os a vi llo sa b. longicornutum g. villosavillosa d. liostracus d . gracilis e. pseudapertum p. wheatleyensishiatus hiatus d . l io st ra cu s cretaceous (pars) lower (pars) upper (pars) jurassic l.l.l.l.l. u.l.u.u.u.u. system series thetys (standard) boreal substage m.u. kimmeridgianvalanginian berriasian tithonian ryazanian volgian barremian hauterivian st ag e cretaceous (pars) lower (pars) upper (pars) jurassic l.l.l. u.l.u.u. system series thetys (standard) boreal substage m.u. kimmeridgianvalanginian berriasian tithonian ryazanian volgian st ag e bernbjerg fm hall bredning group lb wollaston forland gp pb abrr st r 11 0 12 0 13 0 14 0 15 0 16 0 17 0 18 0 19 0 20 0 21 0 22 0 22 5. 7 m 10 20 30 40 50 60 70 80 90 10 0 10m 20 30 40 50 60 70 80 90 10 0 11 0 12 0 13 0 14 0 15 0 16 0 17 0 18 0 19 0 20 0 21 0 22 0 23 0 bernbjerg fmlindemans bugt fm wollaston forland group hall bredning group pb ab h . k oc hi a nd u pp er g . v ill os a vi llo sa lo w er o . c om pl ex up pe r o . c om pl ex r ød ry gg en -1 b ro rs on h al vø -1 https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 33 of 39 geusbulletin.org biostratigraphy. common g. mutabilis characterises the dinoflagellate cyst assemblage in the uppermost p. elegans zone in milne land and ammonite faunal horizon m 25, and p. wheatleyensis zone correlates with the top of this interval (piasecki 1996). the last occurrences of abundant o. patulum and s. clavellii occur in the upper part of the interval, followed by a change in composition of the dinoflagellate cyst assemblage in overlying strata. age. early volgian, late jurassic. organic matter. abundant amorphous kerogen together with dinoflagellate cysts dominates; terrestrial plant material is less significant, occurring mainly as sporomorphs but also some black woody material. the presence of dinoflagellate cysts indicates a marine environment, and the dominance of aom is suggestive of oxygen-deficient bottom conditions. 6.5 unconformity (45.5 m) the occurrence of the d. liostracus – e. pseudapertum zones immediately overlying the p. wheatleyensis zone defines a significant stratigraphic break. the unconformity corresponds to the p. huddlestoni – p. communis zones, spanning the upper lower volgian to lower middle volgian. biostratigraphically, the unconformity is located between 42.09 and 46.05 m and placed at the lithostratigraphic boundary at 45.5 m. 6.6 dorsoplanites liostracus – epipallasiceras pseudapertum chronozones undiff. (45.5–37.50 m) fossils. the zone is based on the dinoflagellate cyst record and correlated with the succession in the rødryggen-1 core from the first appearance of lagenorhytis delicatula (42.09 m) to the top of leptodinium subtile (38.30 m). cribroperidinium spp. dominates the assemblage, and rhynchodiniopsis spp. is common. the lower boundary, however, is placed at the inferred unconformity at the lithostratigraphic boundary at 45.5 m. the assemblage is relatively diverse and abundant. biostratigraphy. the unusual appearance of lagenorhytis delicatula as stratigraphically low as the middle volgian strata is recorded in the rødryggen-1 core (this study). in contrast, the last occurrence of leptodinium subtile is a well-known stratigraphic marker in north-west europe, where it has its highest occurrence in the progalbanites albani (ammonite) zone (riding & thomas 1992). the upper p. albani zone is correlated with the e. pseudapertum zone in the east greenland ammonite zonation (birkelund et al. 1984). the top occurrence of l. subtile just below the ammonite epipalliceras pseudapertum in the rødryggen-1 core (this study) supports this correlation. characteristic boreal, uppermost jurassic species, e.g. paragonyaulacysta borealis, p. capillosa and the acritarch wallodinium krutzschii, that are common in the lower core interval disappear in this zone and are absent throughout the rest of the overlying uppermost jurassic interval. age. middle volgian, late jurassic. organic matter. the interval contains abundant amorphous kerogen with little terrestrial plant material, mainly sporomorphs and a limited content of dinoflagellate cysts. the presence of dinoflagellate cysts indicates a marine environment, and the dominance of aom is suggestive of oxygen-deficient bottom conditions. 6.7 unconformity (37.50 m) the g. villosa villosa dinoflagellate cyst zone, here equivalent to the h. kochi ammonite chronozone, directly overlies the p. liostracus – e. pseudapertum zones. it marks a major stratigraphic break with a hiatus corresponding to the c. anguinus – p. maynci zones interval, spanning the middle middle volgian to the lower ryazanian (fig. 14). 6.8 gochteodinia villosa villosa zone (neg cr 1) and calcareous nannofossil zone bc5 (37.35–30.52 m) fossils. the palynozone is recognised only in one sample at 37.35 m. the assemblage is poor, low density, but due to hardly any other organic matter in this sample, the whole assemblage is recovered in one slide. the assemblage is characterised by a high diversity, especially compared to the organic-rich samples analysed from deeper levels in the core. the interval is barren of dinoflagellate cysts in the upper part (36.36– 30.52 m). the calcareous nannofossil zone bc5 has its base at 37.35 m, 15 cm above the base of the calcareous albrecht bugt member (fig. 17). the upper boundary is placed at 30.52 m, where the assemblage is characterised by very low diversity without age-diagnostic nannofossils. the assemblage of bc5 consists of common to abundant crucibiscutum salebrosum, watznaueria spp., few biscutum constans, rhagodiscus asper, triquerhabdulus shetlandensis, staurolitithes crux, zeugrhabdotus spp., cretarhabdus spp. and discorhabdus spp. (fig. 17). biostratigraphy. the combined presence of dinoflagellate cysts rotosphaeropsis thule and palaecysta palmula https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 34 of 39 geusbulletin.org indicates a lower cretaceous, mid-ryazanian assemblage. r. thule has its range top in uppermost lower ryazanian, h. kochi zone in the rødryggen-1 core. p. palmula is reported to appear in the top of the h. kochi zone, uppermost lower ryazanian in north-west europe (riding & thomas 1992) and in the rødryggen-1 core (this study), where it appears just above an h. kochi ammonite. the absence of abundant oligosphaeridium complex supports the ryazanian age, since its first appearance is in the lowermost valanginian as reported both in north-west europe (riding & thomas 1992) and in the rødryggen-1 well (this study) as well as in the p. albidum ammonite zone, uppermost ryazanian on store koldewey, northern east greenland (nøhr-hansen et al. 2020). the g. villosa villosa zone thus correlates with the h. kochi (ammonite) zone. the calcareous nannofossil assemblage suggests the presence of the calcareous nannofossil zone bc5 of bown et al. (1998). common t. shetlandensis and the high abundance of c. salebrosum in the lower four samples and the co-occurrence of eiffellithus striatus with t. shetlandensis in the upper three samples further suggest subdivision of the interval and the identification of the lower and upper bc5-zonal intervals, respectively (fig. 17). age. latest early ryazanian – valanginian. the dinoflagellate cysts indicate the base of the interval to be latest early ryazanian. the calcareous nannofossil assemblages, on the other hand, indicate a late valanginian age that is somewhat younger than the palynostratigraphic age derived from the basal sample of the interval. previous studies of the ammonite assemblage collected from outcrops of the albrechts bugt member at the drill site suggested a late ryazanian – late early valanginian age (alsen 2006). a late valanginian age based on the bc5 zone also significantly differs from the bc1 zone and late ryazanian age of the base of the albrechts bugt member in the wollaston forland area (pauly et al. 2012a). a slight revision of stratigraphic ranges of ryazanian– valanginian calcareous nannofossil zones in the boreal scheme was based on sr-isotope and c-isotope stratigraphy (möller et al. 2015). the offset in ages obtained from dinoflagellate cysts and calcareous nannofossils observed is, however, far beyond the smaller adjustments recorded by möller et al. (2015). organic matter. a very little organic content comprising rounded to angular, carbonised grains with some dinoflagellate cysts and hardly any terrestrial plant material. this abrupt change in the nature of the organic fraction in the brorson halvø-1 core reflects a significant hiatus, whereas a corresponding rapid environmental change in the rødryggen-1 well takes place over approximately 1 m of transitional beds in the upper ryazanian (this study). 6.9 calcareous nannofossil zones bc6–bc7? (30.52–22.22 m) fossils. this interval is characterised by the lack of age-diagnostic calcareous nannofossils hampering a precise age assessment. the interval is barren of dinoflagellate cysts. the absence of the calcareous nannofossils t. shetlandensis, eprolithus antiquus and tegumentum octiformis indicates the possible presence of the zones bc6–bc7, and the interval is tentatively assigned to bc6–bc7? (fig. 17). age. the bc6 and bc7 zones are considered early hauterivian in age (bown et al. 1998). organic matter. the interval has a very low organic content of angular to rounded carbonised grains. 6.10 calcareous nannofossil zone bc8 (22.22–9.18 m) fossils. the assemblage contains common to abundant watznaueria spp., r. asper, b. constans, zeugrhabdotus spp. as well as few d. ignotus, eiffellithus striatus, perissocyclus spp. and rare sollasites horticus, c. salebrosum, bukrylithus ambiguus and pentaliths; all are characteristic of the higher rødryggen member. the interval is barren of dinoflagellate cysts. biostratigraphy. the assemblage suggests the presence of the bc8 zone. e. striatus, t. octiformis, a few stradnerlithus silvaradius and rare c. salebrosum in the lower part indicate the presence of the bc8a–b subzone; perissocyclus plethotretus and z. scutula in the upper part indicate the overlying bc8c subzone (fig. 17). age. late early hauterivian – early late hauterivian. in contrast to the conflicting ages obtained for the underlying albrechts bugt member (interval 37.50–30.52 m), the calcareous nannofossil age indicated for this interval, the upper part of the rødryggen member, agrees well with the previously recorded age of the rødryggen member (alsen 2006; alsen & mutterlose 2009; pauly et al. 2012a). organic matter. the interval has a very low organic content of angular to rounded carbonised grains. 6.11 batioladinium longicornutum zone (i), n. kostromiensis palyno subzone (i1) and calcareous nannofossil zone bc9 (9.18–6.16 m) fossils. the base of the zone is placed at the first occurrence of the palynomorphs nelchinopsis kostromiensis, muderongia staurota and m. tetracantha. the index fossil of the subzone is thus present, whereas the index of the https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 35 of 39 geusbulletin.org batioladinium longicornutum longicornutum zone was absent. the dinoflagellate cyst assemblage is characterised by oligosphaeridium asterigerum, oligosphaeridium complex, circulodinium distinctum, clesistosphaeridium aciculare, muderongia staurota and m. tetracantha. oligosphaeridium asterigerum and circulodinium distinctum dominate the assemblage (supplementary data file 2). the assemblage is relatively poor in the lower part but becomes diverse and abundant upwards. the calcareous nannofossil assemblage is characterised by common to abundant watznaueria spp., r. asper, b. constans and zeugrhabdotus spp. it contains few d. ignotus, eiffellithus striatus and perissocyclus spp., and rare sollasites horticus, c. salebrosum, bukrylithus ambiguus and pentaliths. biostratigraphy. muderongia staurota and m. tetracantha have their first occurrences in the upper hauterivian s. gottschei ammonite zone according to costa & davey (1992) and duxbury (2001). the highest occurrence of nelchinopsis kostromiensis is in the upper hauterivian? – lower barremian nelchinopsis kostromiensis subzone (i1) in nøhr-hansen (1993) and nøhr-hansen et al. (2020). the presence of the calcareous nannofossils e. striatus, p. plethotretus, p. tayloriae, t. octiformis, s. colligata and z. scutula combined with the absence of c. salebrosum, s. silvaradius, tegulalithus septentrionalis and clepsilithus maculosus suggest the bc9 zone (fig. 17). age. late hauterivian to early barremian based on palynostratigraphy. calcareous nannofossils indicate a late hauterivian age. organic matter. the interval is characterised by a very low organic content that is dominated by black angular to rounded carbonised grains with few dinoflagellate cysts; the latter testify to a marine depositional environment. remarks. it is noteworthy that this interval, which is in the lower part of the stratumbjerg formation, yields calcareous nannofossils despite a return to black mudstone deposition. calcareous nannofossils are thus not restricted to the marls of the underlying albrechts bugt and rødryggen members. despite a sharp boundary between the palnatokes bjerg formation and the stratumbjerg formation observed in the field (figs 5, 6), there appears to be a gradual transition between the two units in terms of palaeoecology and calcareous nannoplankton production. 6.12 batioladinium longicornutum zone (i), p. anaphrissum subzone (i2) (6.16–0 m) fossils. the subzone is based on the dinoflagellate cyst record in one sample (fig. 16; supplementary data file 2). the base of the subzone is recognised at the last occurrence of nelchinopsis kostromiensis. the subzone ranges to the last occurrences of pseudoceratium anaphrissum. oligosphaeridium asterigerum, oligosphaeridium complex, circulodinium distinctum, clesistosphaeridium aciculare and muderongia tetracantha dominate the assemblage. biostratigraphy. the highest occurrence of nelchinopsis kostromiensis is in upper hauterivian? to lower barremian, nelchinopsis kostromiensis subzone (i1) in nøhr-hansen (1993). the highest occurrences of batioladinium longicornutum and hystrichodinium aborispinum are in the upper barremian, whereas the highest occurrence of pseudoceratium anaphrissum is in the upper part of the lower barremian pseudoceratium anaphrissum subzone (i2) in nøhr-hansen (1993) and nøhr-hansen et al. (2020). age. early barremian, early cretaceous. organic matter. the interval has a very low organic content dominated by black, angular, carbonised grains with few dinoflagellate cysts, the latter of which indicate a marine depositional environment. 7. discussion 7.1 the jurassic–cretaceous boundary in (north-east) greenland the jurassic system is characterised by a highly detailed subdivision by means of ammonite zones, which form the standard biostratigraphic and chronostratigraphic framework. ammonite provincialism in some intervals hampers direct correlation to the standard ammonite zonation (established in north-west europe), and secondary standards are then established locally. difficulties in ammonite correlation are particularly pronounced around the jurassic–cretaceous boundary leading to separate chronostratigraphic divisions even at the stage level in separate faunal provinces. the ammonite zonation for most of the upper jurassic in east greenland is mainly based on the ammonite successions in milne land and jameson land (callomon & birkelund 1982; birkelund et al. 1984; birkelund & callomon 1985). the oxfordian ammonite zonation used in east greenland belongs to the boreal and subboreal faunal provinces (sykes & callomon 1979; birkelund et al. 1984; zeiss 2003). less provincialism during the kimmeridgian allows the ammonite succession in greenland to be referred to the standard ammonite zonation of north-west europe (birkelund & callomon 1985; zeiss 2003). provincialism increased progressively during the latest jurassic. the lower volgian ammonite zonation in greenland is adopted from england, whereas a separate boreal zonation was https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 36 of 39 geusbulletin.org established for the middle volgian in greenland, allowing only few correlation levels to the subboreal zonation of england (callomon & birkelund 1982; zeiss 2003). the milne land area is characterised by a gap in the sedimentary succession between the middle volgian and the lowermost cretaceous, whereas the upper volgian in southern jameson land is documented by scattered biostratigraphic data. further to the north, a more complete jurassic–cretaceous boundary succession exists in the wollaston forland – kuhn ø area. the thick, coarse-grained succession represents synrift deposition and contains only sporadic levels with ammonites. the ammonite succession was documented by maync (1949), donovan (1964) and surlyk (1978) providing a zonation that covers the middle volgian to ryazanian. only one zone (praechaetaites tenuicostatus zone) has been assigned to the upper volgian; this has since been referred to the uppermost middle volgian (rogov 2020). however, the presence of subcraspedites (swinnertonia) from kuhn ø (casey 1973) and a record of this genus in the present study indicate the presence of the subboreal subcraspedites primitivus zone in greenland, providing a correlation of the base upper volgian from greenland to north-west europe. the ammonite zonation of the upper jurassic to lower cretaceous interval of relevance to the present study is shown in fig. 14. note the usage of boreal stage names and the position of the boreal upper jurassic – lower cretaceous boundary at the volgian–ryazanian boundary. the jurassic–cretaceous boundary in the standard chronostratigraphy established in the tethyan realm is placed at the tithonian–berriasian boundary. that level probably corresponds to the middle–upper volgian boundary, implying that the uppermost boreal jurassic is to be considered lowermost cretaceous in an international sense. the definition of the base of the cretaceous system in terms of a gssp still remains unresolved. it is agreed that the gssp should be located in the tethyan realm, but it remains to be determined what proxies (e.g. calpionelid microfossils and magnetostratigraphy) are most appropriate for defining the boundary. once formal recognition is achieved, correlation with the boreal and arctic regions can be undertaken. the sedimentary cores recovered at rødryggen-1 and brorson halvø-1 may thus become important reference sections for the jurassic–cretaceous boundary in the boreal realm as they record a succession across the boundary. in this regard, rødryggen-1 is particularly significant with its biostratigraphically nearly unbroken section (fig. 19). 7.2 basin evolution deposition of the palnatokes bjerg formation marked an abrupt end to the late jurassic dysoxic-anoxic–euxinic basin environment and black, organic-rich mudstone deposition (the bernbjerg formation and the storsletten member of the lindemans bugt formation). the palnatokes bjerg formation was deposited in the final, waning phase of rifting, during which the tilted and rotated jurassic fault blocks possibly underwent a last phase of faulting, which led to an even finer subdivision of the blocks. the differences in lithology and thickness observed in the albrechts bugt and rødryggen members reflect strong local control on deposition. for example, the albrechts bugt member at the rødryggen-1 and brorson halvø-1 well sites at stratumbjerg and at perisphinctes ravine (kuhn ø) is characterised by highly condensed, relatively carbonate-rich mudstones that were deposited on local submarine highs (surlyk 1978; alsen 2006). the carbonate content is mainly derived from calcareous nannofossils (pauly et al. 2012a). the member is markedly thicker, less condensed and carbonate poor at, for example, kuhnpasset and niesen (surlyk 1978), reflecting greater accommodation space and higher sedimentation rates in deeper parts of the basin. this probably reflects a change from half-graben basins on a rotated, westerly-tilted block during the middle jurassic to volgian to greater fragmentation in the late ryazanian to hauterivian, when deposition occurred in smaller horst and graben systems. the light grey and red sediments of the condensed albrechts bugt and rødryggen members, deposited on submarine highs, reflect well-oxygenated conditions related to a period of cold climate (alsen 2006). oceanographic changes driven by deep-water formation in the northern proto-north atlantic or in the boreal arctic sea resulted in a palaeo-gulf stream that allowed for the immigration of several tethyan faunal and floral elements to eastern greenland (ager 1971; alsen 2006; pauly et al. 2012b). time-equivalent units, which are compositionally similar to the calcareous albrechts bugt and rødryggen members, occur scattered in the northern north atlantic and the arctic, for example, on andøya, on the norwegian shelf, in the barents sea and at svalbard (kong karls øya). these units show that oceanic currents ventilated the sea water throughout the greenland–norwegian rift basin and the arctic sea during the late ryazanian–valanginian–hauterivian. the palnatokes bjerg formation is overlain by the stratumbjerg formation, which marks the return to grey, dark mudstone deposition. deposition of the stratumbjerg formation reflects the rising sea level during postrift thermal subsidence. in most places, the base of the formation is in the barremian (bjerager et al. 2020), but at some localities, including the brorson halvø-1 core, the transition from the red mudstones of the rødryggen member to the grey mudstones https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ alsen et al. 2023: geus bulletin 55. 8342. https://doi.org/10.34194/geusb.v55.8342 37 of 39 geusbulletin.org of the stratumbjerg formation occurs in the upper hauterivian (bjerager et al. 2020; nøhr-hansen et al. 2020). 8. conclusions the rødryggen-1 core (234.4 m thick) is assigned to the bernbjerg formation (234.4 (td) – 97 m), the lindemans bugt formation (97–24.40 m) and the palnatokes bjerg formation (24.40–0 m). the lindemans bugt formation is represented by a new member, the storsletten member, which records deposition of dark mudstones on the permpas block. this region is detached from the coarse-grained depositional system of the western fault blocks fringing the main rift basin bounding fault. the age of the cored succession is determined from integrated ammonite, palynofossil and calcareous nannofossil stratigraphy and ranges from the upper kimmeridgian a. eudoxus ammonite zone to the upper valanginian o. complex dinoflagellate cyst zone. there are essentially no biostratigraphic gaps in the succession. the brorson halvø-1 core (225.70 m thick) is assigned to the bernbjerg formation (225.7 (td) – 45.5 m), the lindemans bugt formation (storsletten member; 45.5–37.5 m), the palnatokes bjerg formation, including the albrechts bugt (37.5–30 m) and rødryggen members (30–8.7 m), and the stratumbjerg formation (8.7–0 m). the age of the cored succession is determined from integrated ammonite, palynofossil and calcareous nannofossil stratigraphy and ranges from the late kimmeridgian a. eudoxus ammonite zone to the lower barremian b. longicornutum dinoflagellate cyst zone or p. anaphrisum subzone. the succession is characterised by two major biostratigraphic gaps representing unconformities in the succession that bound the storsletten member. they reflect rotation of the permpas block and uplift of the block crest during the middle volgian rift climax, which resulted in repeated non-deposition in contrast to the deeper basin setting of rødryggen-1. both cores, but in particular, the stratigraphically complete rødryggen-1 core, provide key stratigraphic reference sections for the jurassic–cretaceous boundary interval in greenland, with wider implications in the arctic and north atlantic regions. acknowledgements a. ryge and d. samuelsen prepared the palynofossil slides and calcareous nannofossil slides. j. halskov and m.c. porse did the drafting. data were collected under the auspices of geus’ petroleum geological studies, data and services in east and north-east greenland, a collaboration project between geus and several industry companies (bojesen-koefoed et al. 2014). we acknowledge the useful reviews by s. schneider and c. schröder-adams, and editorial comments from jon r. ineson and jørgen a. bojesen-koefoed. additional information funding statement core drilling and fieldwork were undertaken during the geological survey of denmark and greenland (geus) field and drilling campaign in north-east greenland in 2009 and 2010. the subsequent study was funded by a consortium of industry energy companies and the geus. author contributions pal: concept. writing – original draft (lead), ammonites and stratigraphy, lithostratigraphy, discussion, editing spi: concept. writing – original draft (second lead), palynostratigraphy jurassic interval hnh: writing – palynostratigraphy cretaceous interval es: writing – calcareous nannofossil stratigraphy rødryggen-1 core spa: writing – calcareous nannofossil stratigraphy brorson-1 core jh: writing – lithostratigraphy competing interests the authors declare no competing interests. additional files two figures are available as supplementary files at https://doi. org/10.22008/fk2/ktpo5r. supplementary data file 1: a full range chart for palynomorphs in the rødryggen-1 core is provided to supplement fig. 7. supplementary data file 2: a full range chart for palynomorphs in the brorson halvø-1 core is provided to supplement fig. 16. references ager, d.v. 1971: space and time in brachiopod history. in: middlemiss, f.a., rawson, p.f. & newall, g. 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(eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 75–114. https://doi.org/10.34194/geusb.v1.4649 https://doi.org/10.34194/geusb.v55.8342 http://www.geusbulletin.org/ https://doi.org/10.1017/s0016756819001043 https://doi.org/10.34194/geusb.v55.8309 https://doi.org/10.1016/j.cretres.2011.11.011 https://doi.org/10.1016/j.cretres.2011.11.011 https://doi.org/10.1016/j.marmicro.2012.04.004 https://doi.org/10.1016/j.marmicro.2012.04.004 https://doi.org/10.37570/bgsd-2013-61-02 https://doi.org/10.34194/geusb.v1.4650 https://doi.org/10.2113/0240021 https://doi.org/10.1134/s0869593810050047 https://doi.org/10.1134/s0031030120100068 https://doi.org/10.1007/s11430-009-0182-0 https://doi.org/10.1007/s11430-009-0182-0 https://doi.org/10.1134/s0869593811010047 https://doi.org/10.34194/bullggu.v123.6665 https://doi.org/10.34194/bullggu.v128.6670 https://doi.org/10.34194/bullggu.v128.6670 https://doi.org/10.1306/0c9b296b-1710-11d7-8645000102c1865d https://doi.org/10.1306/0c9b296b-1710-11d7-8645000102c1865d https://doi.org/10.34194/geusb.v1.4674 https://doi.org/10.1016/0037-0738(83)90083-0 https://doi.org/10.1016/0037-0738(83)90083-0 https://doi.org/10.34194/bullggu.v105.6646 https://doi.org/10.34194/geusb.v46.6521 https://doi.org/10.1111/j.1502-3931.1976.tb00984.x https://doi.org/10.7306/vj.18.5 https://doi.org/10.1134/s0869593808040059 https://doi.org/10.34194/geusb.v1.4649 stratigraphy of the upper jurassic to lowermost cretaceous in the rødryggen-1 and brorson halvø-1 boreholes, wollaston forland, north-east greenland 1. introduction 2. previous stratigraphic studies of the upper jurassic and lower cretaceous in wollaston forland 3. geological framework and lithostratigraphy 3.1 bernbjerg formation (vardekløft group) 3.2 lindemans bugt formation (wollaston forland group) 3.2.1 storsletten member 3.3 palnatokes bjerg formation (wollaston forland group) 3.3.1 albrechts bugt member 3.3.2 rødryggen member (only in the brorson halvø-1 core) 3.4 stratumbjerg formation (brorson halvø group; only in brorson halvø-1 core) 4. biostratigraphic methods and approach 4.1 macrofossils 4.2 palynomorphs and calcareous nannofossils 4.3 stratigraphic nomenclature and methodology 4.4 rødryggen-1 core 4.5 brorson halvø-1 core 5. biostratigraphy of the rødryggen-1 core 5.1 aulacostephanus eudoxus chronozone (234.40 m (td) 220.51 m) 5.2 aulacostephanus autissiodorensis chronozone (220.51-150.25 m) 5.3 pectinatites elegans chronozone (150.25-130.26 m) 5.4 (pectinatites scitulus -) pectinatites wheatleyensis chronozones undiff. (130.26-119.70 m) 5.5 pectinatites huddlestoni – pectinatites pectinatus chronozones undiff. (119.70–97.00 m) 5.6 dorsoplanites primus chronozone (and pavlovia iatriensis and pavlovia rugosa chronozones; 97.00–79.74 m) 5.7 pavlovia communis chronozone (79.74-74.20 m) 5.8 dorsoplanites liostracus chronozone (74.20-70.15 m) 5.9 dorsoplanites gracilis chronozone (70.15–59.85 m) 5.10 epipallasiceras pseudapertum chronozone (59.85-54.53 m) 5.11 laugeites groenlandicus chronozone (54.53-50.14 m) 5.12 praechetaites exoticus chronozone (50.14-45.37 m) 5.13 subcraspedites primitivus chronozone (45.37–37.49 m) and lowermost gochteodinia villosa villosa zone (neg cr 1; 37.71–37.49 m) 5.14 praetollia maynci chronozone and lower gochteodinia villosa villosa zone (neg cr 1; 37.49–27.73 m) 5.15 hectoroceras kochi chronozone and lower gochteodinia villosa villosa zone (neg cr 1; 27.73–25.82 m) 5.16 upper gochteodinia villosa villosa zone (neg cr 1; 25.82-17.82 m) 5.17 lower oligosphaeridium complex zone (neg cr 2; 17.82-10.17 m) 5.18 upper oligosphaeridium complex zone (neg cr 2; 10.17-1.10 m) 6. biostratigraphy of the brorson halvø-1 core 6.1 aulacostephanus eudoxus chronozone (225.7 m (td) 195.73 m) 6.2 aulacostephanus autissiodorensis chronozone (195.73-149.83 m) 6.3 pectinatites elegans chronozone (149.83-65.23 m) 6.4 pectinatites wheatleyensis chronozone (65.23-45.5 m) 6.5 unconformity (45.5 m) 6.6 dorsoplanites liostracus epipallasiceras pseudapertum chronozones undiff. (45.5-37.50 m) 6.7 unconformity (37.50 m) 6.8 gochteodinia villosa villosa zone (neg cr 1) and calcareous nannofossil zone bc5 (37.35-30.52 m) 6.9 calcareous nannofossil zones bc6-bc7? (30.52-22.22 m) 6.10 calcareous nannofossil zone bc8 (22.22-9.18 m) 6.11 batioladinium longicornutum zone (i), n. kostromiensis palyno subzone (i1) and calcareous nannofossil zone bc9 (9.18–6.16 m) 6.12 batioladinium longicornutum zone (i), p. anaphrissum subzone (i2) (6.16-0 m) 7. discussion 7.1 the jurassic-cretaceous boundary in (north-east) greenland 7.2 basin evolution 8. conclusions acknowledgements additional information funding statement author contributions competing interests additional files references table table 1 summary of macrofossils recorded in the rødryggen-1 core. figures fig. 1 simplified geological map showing the distribution of permian–cretaceous sedimentary rocks in north-east greenland. position of study area in fig. 2 indicated with a red box. reproduced from bojesen-koefoed et al. (2023a, this volume, fig. 1). fig. 2 geological map of the study area. khb: kuhn ø block. kpb: kuppel block. pb: permpas block. hb: hühnerbjerg block. modified from bojesen-koefoed et al. (2023a, this volume, fig. 1). fig. 3 oxfordian-barremian lithostratigraphy for wollaston forland. l.: lower. m.: middle. u.: upper. fig. 4 sedimentological logs of (a) the rødryggen-1 core, (b) the brorson halvø-1 core and (c) legend to both logs. api cs.: american petroleum institute units. vf: very fine. f: fine. m: medium. c: coarse. vc: very coarse. fig. 5 views of the rødryggen-1 drill site. (a) oblique aerial view, towards the south-east of the rødryggen ridge showing the outcropping units and the position of the rødryggen-1 drill site (red dot) situated on a plateau within the yellowish weathering albrechts bugt member. the badlands area in the background is the upper jurassic bernbjerg formation. (b) view across the storsletten plain towards the east and the rødryggen fig. 6 views of the brorson halvø-1 drill site. (a) oblique aerial view of the brorson halvø-1 drill site (red dot) at the foot of the southern slope of bern plateau, sw brorson halvø. (b) geology around the drill site. pale, yellowish weathering of the albrechts bugt member is seen sharply overlain by the rødryggen member (both palnatokes bjerg formation), overlain by the stratumbjerg formation. a roughly se–nw-oriented doleritic dyke intersects the stratumbjerg formation and a doleritic sill caps the bern plateau (sill; upper right). the drill site was situated on a small plateau in the lower part of the stratumbjerg formation (grey tent on platform). photo: a. ryge. ab: albrechts bugt member. fo: fosdalen formation. rr: rødryggen member. sto: storsletten member. str: stratumbjerg formation. fig. 7 the ranges of key dinoflagellate cyst taxa in the rødryggen-1 core. the dinoflagellate cyst species are stratigraphically arranged according to the succession of their lowest occurrences. a range chart showing all taxa recorded in the core is available in supplementary data file 1. fig. 8 the distribution of calcareous nannofossils in the rødryggen-1 core. fig. 9 the ranges of ammonite taxa recorded in the rødryggen-1 core. fig. 10 selected ammonites recorded in the rødryggen-1 core. a: amoeboceras? sp., level 227.72 m, mguh 34180. b: dorsoplanites sp., level 90.85 m, mguh 34181. c: dorsoplanites primus, level 89.30 m, mguh 34182. d: pavlovia cf. variocostata, level 75.43 m, mguh 34184. e: dorsoplanites aff. liostracus, level 74.20 m, mguh 34185. f: pavlovia cf. corona, level 74.03 m, mguh 34186. g: dorsoplanites jamesoni, level 70.06 m, mguh 34187. h: epipallasiceras cf. pseudapertum, level 55.98 m, mguh 34189. i: laugeites cf. biplicatus, level 53.67 m, mguh 34190. j: laugeites cf. intermedium, level 52.82 m, mguh 34191. k: laugeites cf. planus, level 51.55 m, mguh 34192. l: cf. praechetaites exoticus, level 48.82 m, mguh 34193. m: cf. praechetaites exoticus, level 48.54 m, mguh 34194. n: ammonoidea indet., level 48.37 m, mguh 34195. o: aptychus, level 47.20 m, mguh 34196. p: s. (swinnertonia) cf. subundulatus, level 45.37 m, mguh 34197. q: subcraspedites (swinnertonia) sp. juv., level. 43.21 m, mguh 34198. r: cf. praetollia maynci spath, level 37.49 m, mguh 34199. s: hectoroceras sp., level 27.73 m, mguh 34200. the specimens are stored in the palaeontology type collection at the natural history museum of denmark and each labelled with an mguh number – museum geologica universitas hafniensis. fig. 11 selected biostratigraphically significant dinoflagellate cysts and acritarchs from the rødryggen-1 core. scale bars: 25 μm. a and b: paragonyaulacysta capillosa, sample 234.4 m, slide 2. c and d: paragonyaulacysta borealis, sample 140.25 m, slide 6, and 136.26 m, slide 5. e and f: perisseiasphaeridium, sample 226.22 m, slide 4. g and h: oligosphaeridium patulum, sample 130.26 m, slide 6. i and j: cribroperidinium complexum, low and high focus on the same specimen, sample 234.5 m, slide 5. k and l: trichodinium piaseckii, sample 136.26 m, slide 6. m: senoniasphaera clavellii, sample 90.27 m, slide 5. n: cassiculosphaeridium magna, sample 90.27 m, slide 5. o: wallodinium krutzschii, sample 110.15 m, slide 4. p: muderongia simplex, sample 76.74 m, slide 4. fig. 12 selected biostratigraphically significant dinoflagellate cysts and acritarchs from the rødryggen-1 core. scale bars: 25 μm. a, b and c: lagenorhytis showing variable morphology, sample 54.31 m, slide 7. d: gochteodinia villosa subsp. villosa, sample 56.88 m, slide 4 and 5. e: leptodinium subtile, sample 90.27 m, slide 5. f: perisseiasphaeridium insolitum, sample 59.85 m, slide 4. g, h and i: large pterospermella spp., sample 40.62 m, slide 5, and sample 35.2 m, slide 7. j and k: small pterospermella spp., sample 43.43 m, slide 5 and sample 35.2, slide 7. l and m: isthmocystis distincta, sample 35.02 m, slide 3. n, o, p, q and r: morphological variations of gochteodinia villosa subsp. villosa, (n–q) sample 37.71 m, slide 7, and (r) sample 35.02 m, slide 4. s: circulodinium compta, sample 31.95 m, slide 4. t: scriniodinium pharo, sample 24.1 m, slide 3. fig. 13 selected biostratigraphically significant dinoflagellate cysts and acritarchs from the rødryggen-1 (a–m) and the brorson halvø-1 (n–r) cores. scale bars: 25 μm. a: palaecysta palmula, sample 24.1 m, slide 2. b: rotosphaeropsis thule, sample 27.0 m, slide 7. c: scriniodinium pharo, sample 27.1 m, slide 3. d: paragonyaulacysta borealis, sample 35.2 m, slide 4. e: tubotuberella apatela, sample 25.0 m, slide 9. f: phoberocysta neocomica, sample 13.2 m, slide 2. g: paragonyaulacysta capillosa, sample 21.38 m, slide 2. h: oligosphaeridium complex, sample 10.17 m, slide 2. i: oligosphaeridium complex, sample 17.82 m, slide 2. j: nelchinopsis kostromiensis, sample 10.17 m, slide 2. k: palaecysta palmula, sample 9.00 m, slide 2. l: gochteodinia villosa subsp. villosa, sample 24.1 m, slide 3. m: lagenorhytis delicatula, sample 2.78 m, slide 4. n: batioladinium longicornutum, sample 3.37 m, slide 3. o: muderongia tetracantha, sample 9.18 m, slide 3. p: muderongia staurota, sample 9.18 m, slide 3. q: nelchinopsis kostromiensis, sample 6.16 m, slide 3. r: pseudoceratium anaphrissum, sample 3.37 m, slide 3. fig. 14 kimmeridgian – barremian ammonite zonation for north-east greenland. m 14 to m 47 are faunal horizons recorded in milne land by callomon & birkelund (1982). l.: lower. m.: middle. u.: upper. fig. 15 onychites in rødryggen-1 core. a: mega-onychites, level 89.26 m, mguh 34183. b: micro-onychites, level 62.43 m, mguh 34188. fig. 16 the ranges of key dinoflagellate cyst taxa in the brorson halvø-1 core. the dinoflagellate species are arranged according to the succession of their lowest occurrences. question mark (?): uncertainty in taxonomic identification. r: reworked or redeposited. a range chart showing all taxa recorded in the core is available in supplementary data file 2. h-eb: hauterivian – early barremian. ler: latest early ryazanian. bl-nk: b. longicornutum fig. 17 the distribution of calcareous nannofossils in the brorson halvø-1 core. fig. 18 amoeboceras subkitchini (mguh 34201 from geus 469823) ammonite from upper lower kimmeridgian rasenia cymodoce zone. fig. 19 the correlation between the rødryggen-1 and brorson halvø-1 boreholes; the datum is the base of the palnatokes bjerg formation. note the complete stratigraphy in the rødryggen-1 well compared to the significant hiatuses within the volgian–ryazanian in the brorson halvø-1 core. the zones that are present in the rødryggen-1 well and absent in the brorson halvø-1 well are here shown schematically to onlap the unconformity. although it could be argued that these zones were also represented in the brorson halvø-1 area and subsequently removed by erosion of the crest during tilting of the fault block, the lack of redeposited palynomorphs of these ages in the rødryggen-1 core favours an onlap model. ab: albrechts bugt mb. lb: lindemans bugt formation. pb: palnatokes bjerg fm. rr: rødryggen member. str: stratumbjerg formation. see fig. 4 for legend and grain size. research article gowan 2023: geus bulletin 53. 8355. https://doi.org/10.34194/geusb.v53.8355 1 of 17 paleo sea-level indicators and proxies from greenland in the gapslip database and comparison with modelled sea level from the paleomist ice-sheet reconstruction evan j. gowan1,2* 1department of earth and environmental sciences, kumamoto university, kumamoto, japan; 2kikai institute for coral reef sciences, kagoshima, japan abstract one of the most common ways to assess ice-sheet reconstructions of the past is to evaluate how they impact changes in sea level through glacial isostatic adjustment. paleomist 1.0, a preliminary reconstruction of topography and ice sheets during the past 80 000 years, was created without a rigorous comparison with past sea-level indicators and proxies in greenland. the basal shear stress values for the greenland ice sheet were deduced from the present day ice-sheet configuration, which were used for the entire 80 000 years without modification. the margin chronology was based on previous reconstructions and interpolation between them. as a result, it was not known if the greenland component was representative of its ice-sheet history. in this study, i compile sea–level proxy data into the global archive of paleo sea level indicators and proxies (gapslip) database and use them to evaluate the paleomist 1.0 reconstruction. the last glacial maximum (c. 20 000 years before present) contribution to sea level in paleomist 1.0 is about 3.5 m, intermediate of other reconstructions of the greenland ice sheet. the results of the data-model comparison show that paleomist requires a larger pre-holocene ice volume than it currently has to match the sea-level highstands observed around greenland, especially in southern greenland. some of this mismatch is likely because of the crude 2500 year time step used in the margin reconstruction and the limited last glacial maximum extent. much of the mismatch can also be mitigated if different earth model structures, particularly a thinner lithosphere, are assumed. additional ice in greenland would contribute to increasing the 3–5 m mismatch between the modelled far-field sea level at the last glacial maximum and proxies in paleomist 1.0. *correspondence: evangowan@gmail.com received: 14 jun 2023 revised: 30 aug 2023 accepted: 06 sep 2023 published: 13 nov 2023 keywords: glacial isostatic adjustment, ice sheets, sea level, holocene, model-data comparison abbreviations: gapslip: global archive of paleo sea level indicators and proxies gia: glacial isostatic adjustment grip: greenland ice core project gshhg: global self-consistent, hierarchical, high-resolution geography database holsea: holocene relative sea level kyr bp: thousand years before present lgm: last glacial maximum paleomist: paleo ice sheet margins, ice sheets and topography palsea: paleo constraints on sea level rise sle: sea level equivalent geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 license, permitting free redistribution and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: signe h larsen (geus, denmark) reviewed by: william colgan (geus, denmark); sarah bradley (the university of sheffield, uk) funding: see page 13 competing interests: see page 13 additional files: 13 1. introduction sea-level change is one of the biggest threats to society, caused in part by the retreat of the greenland ice sheet because of global warming (fox-kemper et al. 2021). predicting future changes in sea level is essential to protect coastal infrastructure and human settlements. however, the magnitude and pattern of sea-level changes due to ice-sheet retreat is dependent on the past history of the ice sheet in a process known as glacial isostatic adjustment (gia). gia is the combined result of the balance between water stored in land-based ice and the ocean, time-variable earth deformation caused by variations in the proportion of this storage, and changes to the earth’s gravity from changes in the distribution of mass. since the earth deformation is dependent on the history of iceand water-loading, reconstructions of ice-sheet evolution in the past are needed to forecast the impact of changing sea levels. greenland itself is strongly affected by gia-induced changes in sea level that affect human settlements, and have been implicated, for instance, in the collapse of the norse settlements in western greenland (borreggine et al. 2023). the importance of the greenland ice sheet for projecting future sea-level rise means that it has been the subject of many gia-based reconstructions. some such studies that focus on greenland are highlighted here. tarasov & peltier (2002) tuned a dynamic ice-sheet model, based on a shallow ice https://doi.org/10.34194/geusb.v53.8355 https://orcid.org/0000-0002-0119-9440 mailto:evangowan@gmail.com https://creativecommons.org/licenses/by/4.0/deed.ast gowan 2023: geus bulletin 53. 8355. https://doi.org/10.34194/geusb.v53.8355 2 of 17 www.geusbul let in.org approximation, using holocene relative sea-level observations and temperature and age profiles from the greenland ice core project (grip) ice core. this reconstruction is also used in the global ice-5g (peltier 2004) and ice-6g (peltier et al. 2015) reconstructions. fleming & lambeck (2004) investigated the greenland ice sheet using flowline-based reconstructions originally created by hughes (1981) and hughes et al. (1981), and scaled versions of them. simpson et al. (2009) and lecavalier et  al. (2014) created a reconstruction using a dynamic ice-sheet model, based on a shallow ice approximation, tuned to fit sea-level changes and ice extent at the last glacial maximum (lgm). paleomist 1.0 (paleo ice-sheet margins, ice sheets and topography) is a global ice sheet and topography reconstruction for the past 80 000 years at 2500 year time intervals (gowan et  al. 2021). the ice-sheet component was created using the perfectly plastic ice-sheet model icesheet (gowan et  al. 2016a) using ice-sheet margins that were constrained from chronological, geological, and geomorphological constraints. the icesheet reconstruction was refined through a number of iterations using the gia model selen (spada & stocchi 2007; de boer et al. 2014, 2017). this reconstruction is considered to be preliminary, because of its coarse time step (2500 years), and the fact that it was only evaluated against sea-level indicators and proxies in the centre of the laurentide and eurasian ice sheets. these ice masses were the largest contributors to sea-level variations during the past 80 000 years, therefore smaller contributors such as the greenland ice sheet were not rigorously evaluated. in this paper, i compare deglacial period sea-level indicators and proxies from greenland with the sealevel response calculated from paleomist 1.0. to accomplish this, i compiled the indicators and proxies into an online database called gapslip (global archive of paleo sea level indicators and proxies). my goal is to demonstrate the misfit of the current reconstruction to sealevel indicators and proxies to guide future refinements. further refinements on the geologically constrained icesheet margin and directions of ice flow will be needed to make a more robust reconstruction. it is also necessary to take into account multiple possibilities for the earth rheology structure. 2. sea-level data indicators and proxies 2.1 archives of sea-level data since 2008, the palsea (paleo constraints on sea level rise) project has strived to gather scientists interested table 1 sea-level proxies and indicators across greenland location n marine limiting terrestrial limiting index points references north-eastern greenland kap morris jesup 73 67 6 0 ives et al. (1964); funder (1982); möller et al. (2010); funder et al. (2011a) danmark fjord 30 27 0 3 tauber (1960, 1961, 1964); trautman (1963); ives et al. (1964); funder (1982); håkansson (1982); hjort (1997); funder et al. (2011a); bennike & weidick (2001) frederick e. hyde fjord 16 14 1 1 weidick (1972a, 1973, 1977); funder (1982); landvik et al. (2001) germania land 14 14 0 0 landvik (1994) hochstetter forland 20 12 8 0 weidick (1977); håkansson (1978, 1981); hjort (1979, 1981); björck et al. (1994b) hold with hope 17 16 0 1 hjort & funder (1974); håkansson (1975); weidick (1976, 1977); hjort (1979) independence fjord 12 11 1 0 rubin & alexander (1960); ives et al. (1964); tauber (1966); weidick (1977); funder (1982); funder & abrahamsen (1988); bennike (2002); funder et al. (2011a); j.p. koch fjord 2 2 0 0 landvik et al. (2001) jameson land 17 12 5 0 funder (1971, 1972, 1973, 1978, 1990a); weidick (1972a, 1973, 1974); hjort (1979); ingólfsson et al. (1994); björck et al. (1994a); funder & hansen (1996) kap clarence wyckoff 32 29 0 3 ives et al. (1964); tauber (1964); funder (1982); funder & abrahamsen (1988); funder et al. (2011a) kempe fjord 10 10 0 0 håkansson (1973, 1974, 1976); hjort & funder (1974); weidick (1977); hjort (1979) kong oscar fjord 53 50 0 3 washburn & stuiver (1962); trautman (1963); lasca (1966); håkansson (1972, 1973, 1974, 1975, 1976); hjort & funder (1974); hjort (1979) nansen land 6 6 0 0 weidick (1973); kelly & bennike (1985, 1992); bennike & kelly (1987); landvik et al. (2001) nioghalvfjerdsfjorden 17 17 0 0 bennike & weidick (2001) prinsesse ingeborg halvø 67 63 1 3 ives et al. (1964); funder (1982); håkansson (1987); funder & abrahamsen (1988); bennike (1997); hjort (1997); funder et al. (2011a); strunk et al. (2018) tauber (1961) renland 5 4 1 0 funder (1971); hjort & funder (1974) schuchert dal 97 63 0 34 funder (1972, 1978); weidick (1972a); street (1977); hjort (1979); funder & hansen (1996); hall et al. (2008, 2010) traill ø 19 18 0 1 håkansson (1972, 1973, 1974); hjort (1973, 1979); hjort & funder (1974) young sund 27 8 6 13 weidick (1977); hjort (1979); christiansen et al. (2002); pedersen et al. (2011); bennike & wagner (2012) https://doi.org/10.34194/geusb.v53.8355 http://www.geusbulletin.org gowan 2023: geus bulletin 53. 8355. https://doi.org/10.34194/geusb.v53.8355 3 of 17 www.geusbul let in.org in past sea-level variability to deduce changes in ice  sheet and ocean volume (carlson et  al. 2019). for  the period after the lgm, there has been an effort by palsea to compile sea-level indicators and proxies in a standardised way through the associated holsea (holocene relative sea level) project (khan et  al. 2019). a standardised database for all of greenland has yet to be published. the only data set currently available for greenland that is considered to be compatible with the holsea standards is that for isolation basins (long et al. 2011). for the purposes of assessing reconstructions of the greenland ice sheet, a broader database is required. for this study, i have compiled data for all of greenland (table 1, fig. 1). this compilation is not done with the same level of rigour as a holsea-style database, but rather is an interim product that will be replaced when such a study becomes available. these data are part of the broader gapslip database. the database format contains fewer fields than the holsea data sets, as it is intended to be used in conjunction with comparisons with modelled sea level from gia. the initial construction of this database structure began during previous gia assessment studies (gowan et  al. 2016b, 2021). the current version of gapslip (2.0) features a completely revamped code structure and many bug fixes, and data derived from over 1000 studies. in addition to the greenland data set described in this paper, it also incorporates holsea and holsea-compatible databases for eastern canada (vacchi et  al. 2018), eastern united states (engelhart & horton 2012), the baltic sea (rosentau et al. 2021), north sea (vink et al. 2007), northern russia (baranskaya et al. 2018), southeast asia (mann et al. 2019), australia (larcombe et  al. 1995; belperio et  al. 2002; sloss et  al. 2007; table 1 (continued) sea-level proxies and indicators across greenland location n marine limiting terrestrial limiting index points references north-western greenland bessel fjord 36 3 0 33 weidick (1977); blake (1987a); bennike (2002); mcneely & brennan (2005); glueder et al. (2022) kangerluarsuk (cass fjord) 16 15 1 0 weidick (1977); blake (1987a); bennike (2002); mcneely & brennan (2005) hall land 66 37 0 29 rubin & alexander (1960); england (1985); kelly & bennike (1985, 1992); bennike & kelly (1987); mcneely & mccuaig (1991); mcneely & brennan (2005); glueder et al. (2022) inglefield fjord 10 6 4 0 weidick (1976); fredskild (1985); blake et al. (1996) nordvestø 3 3 0 0 kelly et al. (1999) thule 11 10 0 1 funder (1990b); kelly et al. (1999) tuttulissuaq 1 0 1 0 blake (1987b); fredskild (1985) warming land 4 4 0 0 kelly & bennike (1985, 1992); bennike & kelly (1987) wulff land 3 3 0 0 bennike & kelly (1987); kelly & bennike (1992) south-eastern greenland ammassalik 6 0 2 4 long et al. (2008, 2011) south-western greenland akulliit 24 10 1 13 weidick (1972a, 1974, 1976); jungner (1979); long & roberts (2002); long et al. (2011) alluttoq island 10 0 2 8 long et al. (1999, 2006, 2011) eqalussuit tasiat 5 5 0 0 weidick (1972a, 1974) ikertooq fjord 7 5 0 2 weidick (1972a, 1973); ten brink & weidick (1974); ten brink (1975); van tatenhove et al. (1996) ilulissat 12 2 3 7 weidick (1972a, 1973); long et al. (2006, 2011) itilleq 11 2 0 9 weidick (1972a); long et al. (2009, 2011) kangerluk 9 0 0 9 föged (1989); bennike (1995); rasch (1997); long et al. (2011); souza et al. (2021) kangerlussuaq 34 20 4 10 weidick (1972a, 1972b, 1973); ten brink & weidick (1974); ten brink (1975); van tatenhove et al. (1996); storms et al. (2012); bierman et al. (2018) kannala 33 3 3 27 weidick (1974, 1976); jungner (1979); long et al. (2003, 2011); long & roberts (2003) kapisillit 26 8 17 1 weidick (1968, 1972b, 1975, 1976); fredskild (1973, 1983); mcgovern et al. (1996); weidick et al. (2012); larsen et al. (2014) maniitsoq 5 5 0 0 weidick (1973) nanortalik 24 0 0 24 bennike et al. (2002); sparrenbom et al. (2006b); long et al. (2011) nuuk 44 25 19 0 weidick (1973, 1976); fredskild (1983); berglund (2003); hinnerson-berglund (2004); larsen et al. (2014, 2017) paamiut 10 0 1 9 woodroffe et al. (2014) qaqortoq 30 11 0 19 weidick (1975); bennike et al. (2002); sparrenbom et al. (2006a); fredh (2008); randsalu (2008); long et al. (2011); bierman et al. (2018) qeqertarsuatsiaat 11 11 0 0 weidick (1975); larsen et al. (2014) sisimiut 12 3 0 9 weidick (1972a, 1973); bennike et al. (2011); long et al. (2011) tasiussarsuaq 13 4 9 0 lasher et al. (2020) https://doi.org/10.34194/geusb.v53.8355 http://www.geusbulletin.org gowan 2023: geus bulletin 53. 8355. https://doi.org/10.34194/geusb.v53.8355 4 of 17 www.geusbul let in.org lewis et  al. 2013), and antarctica (briggs & tarasov 2013; ishiwa et  al. 2021). data from the lgm and marine isotope stages 3 and 4 (70 000–27 000 yr bp) are also included (gowan et  al. 2022). the database can be found on github (https://github.com/ evangowan/paleo_sea_level) and will be periodically updated. all radiocarbon dates have been recalibrated using oxcal (bronk ramsey 2009) using the latest calibration curves (heaton et al. 2020; hogg et al. 2020; reimer et al. 2020). the ages in this paper are reported as kyr bp (thousands of years before present, where present is defined as the year 1950). 2.2 data compilation although a data compilation has not been published for greenland, a comprehensive list of studies that have sea-level data is contained in lecavalier et al. (2014). all of the references listed in that paper were checked and the relevant data were included. radiocarbon date lists from laboratories that frequently published data from greenland were also checked. i conducted a literature search to find papers published after lecavalier et al. (2014). in total, there are 1019 data points, which were split into 47 subregions to minimise a possible gradient in the gia signal and to ensure that data cluster geographically (table 1). the names of the subregions were taken from a geographical feature within that area. the data include marine-limiting points (where sea level was located above the elevation of the sample), terrestrial-limiting points (where sea level was located below the elevation of the sample), and sea-level indicators (also called index points) that provide an estimate of past sea-level position within a certain elevation range. note that this compilation does not take into account the possibility of tectonically-induced elevation changes, such as the suspected magnitude >8 earthquake that happened in the early holocene in southern greenland (steffen et  al. 2020). data-model comparison plots for all 47 subregions can be found in the supplementary file s1. an example of the data from kangerlussuaq as plotted in the gapslip database is given in fig. 2. in some cases, the locations of the data were not explicitly stated, and it had to be estimated based on maps and descriptions in the original studies. i used google earth™ to estimate the location in these cases. the location provided by google earth may have uncertainties in the order of 10s of km in places, since the satellite imagery was rectified using gshhg (global self-consistent, hierarchical, high resolution geography database; wessel & smith 1996), which is inaccurate in greenland [for details on the inaccuracy, see here: https://github.com/genericmappingtools/gshhg-gmt/issues/12, see also henriksen et  al. (2000).]. this issue may introduce errors in the model-data comparison, depending on the gradient of the gia response. the plots in this paper use a coastline extracted from the bedmachine greenland version 5 topography dataset (morlighem et  al. 2017, 2022) to avoid this problem. 2.3 vertical interpretation and elevation uncertainties to be a useful constraint, sea-level indicators and proxies must provide context on the past sea-level position relative to present day. geomorphic-based indicators provide an ‘indicative meaning’, in which the relative position of past sea level can be determined based on (a) (b) (c) (d) (e) (f) (g) (h) (i) (j) (k) (l) lgm ice margin present-day ice margin sea-level proxy locations in fig. 5 fig. 1 map showing the locations of the 47 subregions for which there are data in the gapslip database for greenland and the present-day and lgm grounded ice-sheet margin from paleomist. the locations with data-model comparisons shown in fig. 5 are labelled as follows: (a) hall land (b) kap clarence wyckoff (c) germania land (d) young sund (e) schuchert dal (f) ammassalik (g) nanortalik (h) nuuk (i) ikertooq fjord (j) kannala (k) alluttoq island (l) thule. https://doi.org/10.34194/geusb.v53.8355 http://www.geusbulletin.org https://github.com/evangowan/paleo_sea_level https://github.com/evangowan/paleo_sea_level https://github.com/genericmappingtools/gshhg-gmt/issues/12 https://github.com/genericmappingtools/gshhg-gmt/issues/12 gowan 2023: geus bulletin 53. 8355. https://doi.org/10.34194/geusb.v53.8355 5 of 17 www.geusbul let in.org a modern analogue and within a range of uncertainty (rovere et  al. 2016). for instance, a beach deposit will form between the ordinary berm (the upper limit of wave-generated deposition) and the breaking depth of waves. another indicator is an isolation basin, where the transition of the basin from marine to lacustrine deposits will provide the timing for when sea level was positioned at the outlet of a basin (long et al. 2011). in cases where samples are in littoral deposits, they can only be judged as marine limiting and that sea level was above the elevation of the deposits. likewise, archaeological and terrestrial deposits can only indicate that sea level was below the elevation of the sample. most of the data from greenland are terrestrial or marine limiting (table 1). redeposited shells in diamicton and glacial till generally cannot be used as proxies because the geological context cannot be determined. many of the data in the database are from marine shells, with a limited description of the geological context of the deposits. for these data, it is not possible to interpret the water depth, therefore they are included as marine-limiting indicators. where the shells are reported from beach or beach ridge deposits, it is possible to infer a sea-level index point. in these cases, the programme imcalc was used to produce the uncertainty range (lorscheid & rovere 2019). this programme uses models of tidal range and wave heights to infer the uncertainty of the indicative meaning of the deposits. one of the largest sources of data in the database comes from an archive of driftwood from northern and north-eastern greenland reported by funder et  al. (2011a). the driftwood was reported as generally deposited 1–2 m above sea level by storm action, but the authors warned that some samples had likely moved downslope after deposition. the driftwood gives maximum ages for the sea-level position as the trees grew for an unknown time before deposition, and the area where the driftwood was found was undergoing postglacial uplift, i have therefore, conservatively, included these data as marine limiting, after subtracting 2 m from the reported elevation. to compare sea-level indicator and proxy data to modelled sea level, it is necessary to ensure the elevation of the data is reported relative to a known datum, usually mean sea level. the uncertainty on the elevation measurements depend on the technique used. for instance, elevation measurements from differential gps can achieve a precision of less than 10 cm, while elevations derived from topographic maps can be in the order of metres (rovere et al. 2016). in studies where the method used to determine elevation is clearly described, i have used the reported elevation uncertainty. however, the vast majority of the studies incorporated into the 52°w 51°w 50°w 67°n 0 40 80 km kangerlussuaq −40 −20 0 20 40 60 80 100 120 140 e le va tio n (m ) age (kyr bp) # samples: 34 terrestrial limiting index point (>10m) marine limiting index point (≤10m) sea-level proxy type (a) (b) 7 6 5 4 3 2 1 011 10 9 8 fig. 2 paleo sea level and comparison with the reference model at kangerlussuaq subregion. (a) a map of the locations of the data, including a yellow outline that defines the subregion. this location demonstrates the four classes of data, including marine limiting (sea level was above the data point), terrestrial limiting (sea level was above the data point), and index points (sea level was within a bounded elevation range), which has different shades depending on whether or not the uncertainty is less or greater than 10 m. (b) the elevation of the proxy data with uncertainty ranges, and the calculated sea level at the location of each point from the reference paleomist 1.0 model. there is a gradient in the calculated sea level in this area, so multiple calculated sea-level curves are visible. data uncertainties are displayed at 2σ. data references: weidick (1972a, b, 1973); ten brink & weidick (1974); ten brink (1975); van tatenhove et al. (1996); storms et al. (2012); bierman et al. (2018). https://doi.org/10.34194/geusb.v53.8355 http://www.geusbulletin.org gowan 2023: geus bulletin 53. 8355. https://doi.org/10.34194/geusb.v53.8355 6 of 17 www.geusbul let in.org database do not report uncertainties, the datum used for the elevation, the tidal range or details on how the elevation was measured. in these cases, i have applied an uncertainty based on the recommendations in rovere et al. (2016). for reported elevations less than 5 m, an uncertainty of ±1 m is applied. for reported elevations above 5 m, ±20% of the reported elevation is applied, up to a maximum of ±10 m. this level of uncertainty is justified, as it has been reported that some of the reported elevations of legacy greenland data have errors in the order of 10 m (woodroffe et al. 2014). 2.4 age control the vast majority of the age constraints of the sea-level indicators and proxies come from radiocarbon dates (990 points), though there are also some constraints from optically stimulated luminescence (11 points), cosmogenic 10be (15 points), and age models (2 points). all of the data, including calibrated radiocarbon dates, are displayed at 2σ limits. conventional radiocarbon dates are corrected for the isotopic fractionation of carbon by normalising to δ13c = −25‰ relative to the peedee belemnite standard (stuiver & polach 1977). many early radiocarbon laboratories did not follow this standard, and therefore those dates need to be corrected for the fractionation effect before calibration. a large portion of the dates in this database come from marine carbonates that have a value of approximately δ13c = 0‰, which equates to a roughly 400 year offset if not corrected. some laboratories did correct for fractionation, but unconventionally normalised marine carbonates to this value. all the laboratory procedures are documented at https://github.com/ evangowan/radiocarbon_labs. where a fractionation correction was required, i used the estimated values of δ13c listed in stuiver & polach (1977). marine carbonate radiocarbon ages need correction for the offset in age from the atmosphere because of the marine carbon reservoir. the reservoir corrections are derived from the calib reservoir age database (reimer & reimer 2001). for the purposes of correcting data from greenland, i have used two corrections, one for the part of greenland adjacent to baffin bay, and another for the rest of greenland (table 2). 3. model-data comparison 3.1 the paleomist reconstruction of greenland paleomist 1.0 is a preliminary ice-sheet and topography reconstruction for the past 80 000 years at 2500 year resolution (gowan et al. 2021). the goal of this reconstruction was to create a generalised depiction of ice-sheet evolution over this period based on geological and geophysical evidence from the core areas of the north american and eurasian ice sheets, which contributed to the majority of sea-level changes during the last glacial cycle. the ice sheets were constructed using the plastic ice-sheet model icesheet (gowan et al. 2016a) assuming equilibrium conditions. in the most basic version of this model (i.e. without variations in base topography or shear stress), the change in ice-surface elevation, e, at a distance, s, along a flowline is related to the basal shear stress τo through the following equation (cuffey & paterson 2010): de ds i gh 0 ρ = τ the density of ice is ρi and g is the gravity at the surface of the earth. in this formulation, the ice-sheet surface profile is approximated as a parabolic shape. the primary variables in the model were the ice margin and basal shear stress, which control the steepness of the ice-surface profile. the ice margins were largely based on previously published margin reconstructions or geological evidence (e.g. dyke 2004; raised consortium et  al. 2014; hughes et  al. 2016; dalton et  al. 2019). the initial basal shear stress values for the paleo ice sheets in europe and north america were parameterised based on topographic and surficial geological considerations. these values were further adjusted to improve the misfit between the modelled sea level and geological evidence of sea-level change in the core regions of the north american and eurasian ice sheets (baranskaya et  al. 2018; vacchi et  al. 2018; rosentau et al. 2021). in general, it was set so the shear stress values decrease during deglaciation, as the ice sheets likely thinned before the margin retreated. the base topography used for the reconstruction was rtopo-2 (schaffer table 2 reservoir age used to correct marine carbonates location reservoir age calib map number1 references western greenland including baffin bay, davis strait and nares strait 39 ± 107 9, 10, 11, 34, 35, 36, 37, 38, 39, 40, 665, 666, 721, 724, 725, 726, 727, 728, 729, 730, 782, 786, 787, 788, 789, 986, 987, 988, 989, 990, 2062 olsson (1980); mörner & funder (1990); mcneely et al. (2006); coulthard et al. (2010); dyke et al. (2019) eastern and northern greenland −51 ± 71 21, 22, 23, 25, 26, 27, 28, 29, 30, 667, 669, 670, 671, 791 håkansson (1973); tauber & funder (1975); olsson (1980) 1map number refers to the ‘mapno’ field in the calib reservoir age database (https://calib.org/marine/). https://doi.org/10.34194/geusb.v53.8355 http://www.geusbulletin.org https://github.com/evangowan/radiocarbon_labs https://github.com/evangowan/radiocarbon_labs https://calib.org/marine/ gowan 2023: geus bulletin 53. 8355. https://doi.org/10.34194/geusb.v53.8355 7 of 17 www.geusbul let in.org et  al. 2016). the reconstructed ice sheets were calculated on a 5 km resolution grid. changes in sea level and topography were calculated using the gia model selen (spada & stocchi 2007; de boer et al. 2014). the rheology model used for paleomist is a three layer, spherically concentric earth (i.e. a 1d) model with an 120 km thick elastic lithosphere, an upper mantle viscosity of 4 × 1020 pa·s and a lower mantle viscosity of 4 × 1022 pa·s. in reconstructing the ice sheet, the topography and sea-level changes were iterated several times to account for changes in the ice loading. since the greenland ice sheet contributed only a small amount to sea-level change during the last glacial cycle, it was not a focus of investigation for paleomist 1.0. as a result, the margin history and shear stress values were not scrutinised to the same extent as for the north american and eurasian ice sheets. the shear stress values were held constant through the entire time period. the margins were drawn through interpolation between the present margin and the inferred lgm margin through interpolation, with minimal considerations of the impacts of ice-sheet dynamics and topography. the basal shear stress is primarily related to the topographic roughness and basal geology. in greenland, the basal geology is completely unconstrained, so i divided the regions based only on topography (gowan et al. 2016a). the domain boundaries of equal shear stress are based on the locations of fjords, mountain ranges and relatively flat areas in the interior of greenland (fig. 3). the value in each subregion was then tuned to reproduce the ice thickness of the modern ice-sheet configuration. in currently deglaciated coastal shelf regions where the ice sheet interacted with the ocean, the basal shear stress is assumed to be a low value because of the introduction of buoyancy forces and the presence of deforming sediments at the base. the value in each subregion was then tuned to reproduce the  ice thickness of the modern ice-sheet configuration. the shear stress values were held to be constant in the reconstruction, though thinning or reduction in elevation of the central parts of the ice sheet during the holocene (vinther et al. 2009; lecavalier et al. 2013) would imply a reduction in shear stress in the absence of large-scale margin retreat. in general, the shear stress values are relatively high (>100 kpa) around the edges of greenland, where there is mountainous topography. it is lower (<100 kpa) in the centre where the topography is flatter and the ice sheet – surface elevation gradient becomes limited. this is related to the impact on ice-sheet dynamics of the mountains around the edge of greenland that impede ice flow from central greenland (cuffey & paterson 2010). the continental shelf areas are set to have a low nominal shear stress (<10 kpa). the low shear stress is expected because of the interactions with the ocean and the fact that the ice sheet would be underlain with unconsolidated sediments that would encourage ice flow. using the present-day basal shear stress values may cause the ice thickness in the interior of the greenland ice sheet to be overestimated during the glacial period. the core of the ice sheet may have been thinner than at present because of dynamic effects of softer ice from the glacial period and lower accumulation rates (reeh 1985; cuffey & clow 1997), though it may not be possible to quantify this (lecavalier et al. 2013). in paleomist 1.0, the increase in ice thickness in the centre of the greenland ice sheet at the lgm varies between 150 and 300 m. if the interior of the ice sheet was thinner than at present during the lgm, it would increase the potential maximum sea-level highstand in coastal regions because of a reduction of forebulge effects. the margin reconstructions for greenland during the past 80 000 years was based on a number of in ferences basal shear stress (kpa) 0 40 80 120 160 200 fig. 3 basal shear stress values used to reconstruct the greenland ice sheet. https://doi.org/10.34194/geusb.v53.8355 http://www.geusbulletin.org gowan 2023: geus bulletin 53. 8355. https://doi.org/10.34194/geusb.v53.8355 8 of 17 www.geusbul let in.org from geological data and previously published reconstructions. the modern margin was extracted from the rtopo-2 data set; schaffer et al. (2016), which defines the grounded part of the greenland ice sheet. this was also used as the margin at 7.5 kyr bp. for the 5 kyr bp time step, the margin in south-western greenland was set to retreat about 40 km from the present-day margin, based on evidence of retreat between 10 to 80 km from its current extent in the mid-holocene (funder et al. 2011b). the 2.5 kyr bp margin was set to be intermediate of the 5 kyr bp and modern margins. the margins from 10 to 17.5 kyr bp were derived from the reconstructions by dyke (2004). it appears that these margins were drawn using the inaccurate gshhg coastlines mentioned in section 2.2, therefore the reconstructed ice sheet will be in error in north greenland. the lgm extent is based on the reconstruction presented by funder et al. (2011b), and this is also used as the basis for the margin location to 30 kyr bp. the margin reconstructions for 77.5 to 32.5 yr bp were set to be intermediate of the lgm and present-day margin, with fluctuations to coincide with the timing of heinrich events (andrews & voelker 2018), and geological constraints reported in a number of studies (alley et al. 2010; funder et al. 2011b; simon et al. 2014; larsen et al. 2018). the marine isotope stage 4 maximum extent was set to 60 kyr bp, with an extent of 25 km landward from the lgm margin. the margin at 80 kyr bp is set to be the same as present. figure 4 shows the thickness and volume changes of the greenland ice sheet since the lgm. the ice volume is reported as sea level equivalent (sle), which is the ice volume converted to an equivalent amount of ocean-water volume, and divided by the modern area of the ocean fig. 4 difference in ice thickness from the present-day greenland ice sheet in paleomist 1.0, reported as sea-level equivalents (sle) at various time slices. the dark green line is the location of the (grounded) ice-sheet margin. time slices are as follows: (a) 2.5 kyr bp. (b) 5 kyr bp. (c) 7.5 kyr bp. (d) 10 kyr bp. (e) 12.5 kyr bp. (f) 15 kyr bp. (g) 17.5 kyr bp. (h) 20 kyr bp. (a) 2500 yr bp –0.1 m sle (b) 5000 yr bp –0.4 m sle (c) 7500 yr bp 0.0 m sle (d) 10 000 yr bp 2.2 m sle (e) 12 500 yr bp 2.8 m sle (f) 15 000 yr bp 3.4 m sle (g) 17 500 yr bp 3.5 m sle (h) 20 000 yr bp 3.5 m sle ice thickness difference from present (m) –2000 2000150010005000–500–1000–1500 https://doi.org/10.34194/geusb.v53.8355 http://www.geusbulletin.org gowan 2023: geus bulletin 53. 8355. https://doi.org/10.34194/geusb.v53.8355 9 of 17 www.geusbul let in.org (361 x 106 km2). the sle calculation also subtracts the modern ice volume and water volume on the continental shelf. at the lgm, the greenland ice sheet contributed to about 3.5 m of sle ice volume to global ice volume, less than 4% of the total excess ice volume in paleomist 1.0. this value is maintained until the time slice at 12.5 kyr bp, when the ice volume is reduced by 0.6 m sle. at 7.5 kyr bp, the ice margin is set to be the same as present, thus there is essentially no difference in ice volume from the present. the reduced extent margin in western greenland caused a reduction in ice volume of 0.4 m sle. most of the additional ice in the reconstruction at the lgm is located in areas that are currently ice free, with only limited ice thickness gain (<500 m) in much of the interior of the ice sheet. 3.2 comparison of calculated sea level with proxies and indicators the calculated sea level and data for selected locations found in fig. 1 are shown in fig. 5. although the paleomist 1.0 reconstruction has 2500 year time steps, the sea level is calculated by linearly interpolating the ice load to 500 year time steps. this is done to avoid the overestimation of loading caused by the fact that selen treats the load as a heaviside function (i.e. constant ice volume between two time steps). some of the locations contain data from a broad area where the gradient in the gia signal is large. therefore, the calculated sea level after deglaciation within a single location might have a large variations depending on the proximity to the centre of the ice load. this can be seen in the locations with multiple calculated sea-level curves on the figures. overall, the calculated sea level from paleomist 1.0 fails to achieve the high relative sea-level values implied by the sea-level proxy and indicator data for the early holocene. the only locations that come close are in the vicinity of the nares strait, such as hall land. the relatively good fit there is likely a consequence of the fact that the neighbouring innuitian ice sheet was tuned to −80 −40 0 40 80 120 160 e le va tio n (m ) (a) hall land (b) kap clarence wyckoff (c) germania land (d) young sund −80 −40 0 40 80 120 160 e le va tio n (m ) (e) schuchert dal (f) ammassalik (g) nanortalik (h) nuuk −80 −40 0 40 80 120 160 e le va tio n (m ) 024681012 age (kyr bp) (i) ikertooq fjord 024681012 age (kyr bp) (j) kannala 024681012 age (kyr bp) (k) alluttoq island 024681012 age (kyr bp) (l) thule marine limiting terrestrial limiting indicator (≤10m) indicator (>10m) calculated sea level fig. 5 plots showing sea-level indicators and proxies for selected subregions around greenland, and the calculated sea-level curves from paleomist 1.0. a darker shade of green is used for sea-level indicators that have an uncertainty range less than 10 m to emphasise their quality. since the locations of the data often cover a broad area, there can be a gradient in the sea-level response, and so multiple calculated curves are shown. locations in fig. 1 for the following subregions: (a) hall land (b) kap clarence wyckoff (c) germania land (d) young sund (e) schuchert dal (f) ammassalik (g) nanortalik (h) nuuk (i) ikertooq fjord (j) kannala (k) alluttoq island (l) thule. https://doi.org/10.34194/geusb.v53.8355 http://www.geusbulletin.org gowan 2023: geus bulletin 53. 8355. https://doi.org/10.34194/geusb.v53.8355 10 of 17 www.geusbul let in.org fit sea-level data, although for a different earth rheology model than that used in paleomist 1.0 (khosravi 2017). this misfit is particularly pronounced in the southern and western parts of greenland, such as nanortalik and kannala. many of these sites have tightly constrained sea level histories from isolation basin studies (long et  al. 2011), therefore this misfit demonstrates a deficiency in the model. 4 discussion 4.1 exploring potential solutions to the mismatch one of the possible reasons that the calculated sea level was unable to match observations is that the earth rheology structure used for paleomist may be inappropriate. this would be unsurprising, since the value for lithospheric thickness (120 km) is considered to be appropriate in stable precambrian cratons, where much of the laurentide ice sheet was located. though the core of greenland is predominantly precambrian (henriksen et al. 2000), it is also affected by the cenozoic passage of the iceland hot spot (rogozhina et al. 2016). even so, the earth model used in paleomist is similar to the optimal model for eastern greenland found by simpson et al. (2009; 120 km lithosphere, 3 × 1020 pa·s upper mantle and 5 × 1022) and within the range of optimal models found by (lecavalier et al. 2014). the later study found that there was only limited sensitivity to lower mantle viscosity, a result that is consistent with my own analysis, with the exception of northern greenland (see gowan 2023a). milne et al. (2018) found that variations in lithospheric thickness relative to a uniform value of 120 km may be responsible for over 20 m of the observed highstand at the start of the holocene in some areas. the response to deglaciation may also be influenced by time-variable (transient) viscosity of the upper mantle (paxman et al. 2023). the other main possibility is that the history of icesheet volume is inappropriate. the maximum excess ice volume in paleomist at the lgm is 3.5 m sle (fig. 4), which is less than the 4.6 m sle value estimated by simpson et al. (2009) and 4.7 m sle estimated by lecavalier et  al. (2014). however, it is more than the 3.1 m sle estimated by fleming & lambeck (2004) and the 1.9 m sle estimated from tarasov & peltier (2002). the 2500 time step used for the margin history may also fail to capture the precise timing of the retreat of the ice sheet, which was predicted by lecavalier et al. (2014) to have largely happened between 12 and 10 kyr bp. the paleomist model may initiate ice-sheet retreat too early if this is correct, which would decrease the potential sea-level highstand at 10 kyr bp. the model by lecavalier et al. (2014) depicts the western greenland ice sheet as extending to the shelf edge, in contrast to the more restricted extent that was used in paleomist from funder et al. (2011b). if this is correct, there would be a greater perturbation of the upper mantle response as the excess volume would be spread over a larger area, which would result in a larger highstand. to address some of these possibilities, i have run a number of additional earth and ice models, the full results of which can be viewed in the gapslip-paleomist model comparison reports (gowan 2023a). to highlight some of the possibilities, i have selected two different earth models for comparison, one with a lithospheric thickness of 60 km rather than 120 km, and another using the vm5a viscosity model that is used by peltier et al. (2015; fig. 6). a 60 km thick lithosphere has been inferred from parts of greenland affected by the iceland hot spot from the modelling of present-day uplift rates (khan et al. 2016). the model used in fig. 6 has a slightly different upper and lower mantle viscosity than used by khan et al. (2016), but in the interests of assessing the impact of lithospheric thickness changes, i have not changed them to match. the vm5a model has a 60 km thick lithosphere, a 40 km thick layer below the lithosphere with a viscosity of 1 × 1022 pa·s, an upper mantle viscosity of 5 × 1020 pa·s and a lower mantle visosity (between 660 and 1160 km depth) of 1.6 × 1021 pa·s. the rest of the lower mantle has a viscosity of 3.2 × 1021 pa·s. i have also selected an alternative ice-sheet model where the basal shear stress values for greenland have been increased by 20 kpa prior to 10 kyr bp. this has the effect of increasing the lgm ice volume to 4.8 m sle, which is closer to the models by simpson et al. (2009) and lecavalier et  al. (2014). the four locations shown in fig. 6 represent the different parts of northern (kap clarence wyckoff), eastern (schuchert dal), southern (nanortalik) and western (kannala) greenland. the results show that at least for northern, eastern and western greenland, an improved fit can be achieved by reducing the lithospheric thickness, or using the more complex structure of vm5a, without modifications to the ice-volume history. it is possible that the 120 km lithosphere thickness is inappropriate for most of greenland. the improved fit from the vm5a could be the result of having the thin high-viscosity layer under the lithosphere, which delays the rebound after melting compared to the elastic rheology. the weaker lower mantle may also change the position and increase the rate of collapse of the forebulge of the north american ice sheets. for these regions, increasing the ice thickness only has a relatively small improvement on matching the calculated sea level to the data, at least if the ice-margin history is unchanged. the match in southern greenland is not substantially improved by either changing the ice thickness or earth structure, suggesting that substantial https://doi.org/10.34194/geusb.v53.8355 http://www.geusbulletin.org gowan 2023: geus bulletin 53. 8355. https://doi.org/10.34194/geusb.v53.8355 11 of 17 www.geusbul let in.org revisions on the basal shear stress and ice-margin history are required to fit these data. this misfit is not surprising since the paleo ice thickness is not substantially different from present in southern greenland (fig. 4). in all cases, the fit may be improved by including a more detailed history of margin retreat such as the recently released paleogris margin reconstruction (leger et  al. 2023). a final possibility is that the resolution of the gia modelling needs to be increased. the paleomist reconstruction used in selen is composed of disc elements with a radius of approximately 34 km and a spherical harmonic expansion of 256 degrees. these limits were considered appropriate given the preliminary nature of the reconstruction, and the computational expense if the resolution was increased further. if the deep fjords around the margin of greenland bias the elevation of the elements downwards, the programme may interpret the ice as floating and it will not contribute to loading. this bias could be mitigated −80 −40 0 40 80 120 160 e le va tio n (m ) (a) kap clarence wyckoff (default) (b) schuchert dal (default) (c) nanortalik (default) (d) kannala (default) −80 −40 0 40 80 120 160 e le va tio n (m ) (e) kap clarence wyckoff (60 km lithosphere) (f) schuchert dal (60 km lithosphere) (g) nanortalik (60 km lithosphere) (h) kannala (60 km lithosphere) −80 −40 0 40 80 120 160 e le va tio n (m ) (i) kap clarence wyckoff (vm5a) (j) schuchert dal (vm5a) (k) nanortalik (vm5a) (l) kannala (vm5a) −80 −40 0 40 80 120 160 e le va tio n (m ) 024681012 age (kyr bp) (m) kap clarence wyckoff (thick ice sheet) 024681012 age (kyr bp) (n) schuchert dal (thick ice sheet) 024681012 age (kyr bp) (o) nanortalik (thick ice sheet) 024681012 age (kyr bp) (p) kannala (thick ice sheet) marine limiting terrestrial limiting indicator (≤10m) indicator (>10m) calculated sea level fig. 6 plots showing sea-level indicators and proxies for selected subregions around greenland and the calculated sea-level curves from paleomist 1.0 using (a–d) the default earth model, (e–h) a 60 km lithosphere thickness rather than 120 km, (i–l) using the vm5a earth model, and (m–p) using a thicker greenland ice sheet where the shear stress values before 10 kyr bp have been increased by 20 kpa. since the locations of the data often cover a broad area within the region, there can be a gradient in the sea-level response, and so there are multiple calculated curves. locations in fig. 1. a darker shade of green is used for sea-level indicators that have an uncertainty range less than 10 m to emphasise their quality. https://doi.org/10.34194/geusb.v53.8355 http://www.geusbulletin.org gowan 2023: geus bulletin 53. 8355. https://doi.org/10.34194/geusb.v53.8355 12 of 17 www.geusbul let in.org by using higher resolution grid elements or by using a median filter to create the topography rather than an average filter or random sampling. this will be considered in future reconstructions. 4.2 contributions to global sea level the contribution to lgm sea level from the greenland ice sheet in paleomist 1.0 is 3.5 m sle and represents 3–4% of the lgm ice volume. this value is intermediate of other gia-based ice-sheet reconstructions (tarasov & peltier 2002; fleming & lambeck 2004; simpson et al. 2009; lecavalier et al. 2014). in gowan et al. (2022), we concluded that the misfit between the calculated sea level and some far-field sea-level indicators (far from the gia effects of the ice sheets) was because of the lack of smaller ice caps and glaciers, changes in land-based water storage and thermal expansion in the reconstruction. from the results of this study, it is possible that additional ice volume from greenland could also contribute to the underestimate of global ice volume at the lgm in paleomist. the results from this study indicate that there is a certain amount of ambiguity in the greenland ice sheet’s contribution to global sea level at the lgm. the paleomist reconstructed ice sheet, though intermediate of other reconstructions, does not provide a great match to the paleo sea-level observations. more ice, through increased basal shear stress or an extended ice margin, could be added to the reconstruction to reconcile some of these observations, but this can also be countered by accounting for lateral changes in lithosphere thickness and upper mantle viscosity. some recently collected geological constraints, such as cosmogenic dates from western greenland (graham et al. 2019; sbarra et al. 2022) and submarine landform features (ó cofaigh et al. 2013), favour larger ice-sheet configurations. the consequence of this ambiguity means that it is difficult to constrain the history of the ice sheet for the purposes of predicting the future of the ice sheet. the momentum caused by past changes in ice-sheet dynamics have impacts on the current dynamics of the ice sheet, and may be delaying changes because of current global warming (yang et  al. 2022). whether having a larger (i.e. >4.5 m sle) or smaller (i.e. <2 m sle) at the lgm, impacts the current trajectory of ice-sheet retreat should be the subject of further investigation. 4.3 improvements to the sea-level indicator and proxy database the sea-level indicator and proxies presented in this study were compiled in a way that is sufficient to evaluate the fit of calculated sea level. however, the database could be improved to reduce the vertical uncertainties if more details on the survey techniques were found. it may also be possible to determine the elevation with lower uncertainty using modern high-resolution topography data sets. it may be possible to infer the sea-level indicative range for in situ marine molluscs that are currently classified as marine limiting through scrutinising the geological context of the deposits, or by using the depth-range inference techniques proposed by glueder et al. (2022). the gapslip database also excludes marine limit-data that may not be possible to directly date using radiocarbon. marine-limit data are widely available in greenland (dyke et al. 2005) and it may be possible to assign an age based on regional correlations or via cosmogenic dating techniques. since the marine limit can be determined through remote sensing, it may be one way to determine the gia signal in places with few other constraints (e.g. mcmartin et al. 2022). finally, since this database relied on the list of references found in lecavalier et al. (2014) for data published prior to 2014, it is possible that key studies containing additional proxies were omitted. i hope to actively update the database, and welcome additional sea-level proxy data for inclusion as they become available. 5. conclusions this study has presented a new publicly accessible archive of sea-level indicators and proxies for greenland as part of the gapslip database. this archive makes it possible to easily assess calculated sea level from gia models. these data demonstrate that the pre-holocene paleomist 1.0 ice-sheet reconstruction likely requires additional ice volume in greenland, particularly in southern greenland. it is also likely that much of the misfit with sea-level data is attributable to neglecting lateral variations in earth rheology, and that using a thinner lithosphere will produce a better fit to the data without requiring substantially more ice. another possibility is that the lgm margin of the ice sheet was more expansive than inferred. the new sea-level data set could be supplemented with additional data and improved with more refined uncertainty estimates. acknowledgments the figures were generated using generic mapping tools (gmt) version 6.4 (wessel et al. 2019). my gratitude goes to the authors of gmt who fixed a couple of issues i was having, and for clarifying some of the aspects of the greenland coastline reconstruction. some figures in this manuscript make use of the scientific colour maps (crameri et al. 2020). i also thank thomas lorscheid and alessio rovere who gave guidance in my port of their tool imcalc to python, which is included in gapslip. i also thank tomohiko tomita, who has helped me with logistical aspects of my postdoctoral studies, which this paper is part of. i acknowledge palsea, a working group of the international union for quaternary sciences (inqua) and past global changes (pages), which in turn received support from the swiss academy of sciences and the chinese academy of sciences. i thank paolo stocchi for the https://doi.org/10.34194/geusb.v53.8355 http://www.geusbulletin.org gowan 2023: geus bulletin 53. 8355. https://doi.org/10.34194/geusb.v53.8355 13 of 17 www.geusbul let in.org use of his standalone version of selen. lastly, i thank reviewers, william colgan and sarah bradley, for their comments that helped improve the paper. additional information funding statement funding came from an international postdoctoral fellowship of japan society for the promotion of science. author contributions e.j.g compiled the sea level proxy data, performed the analysis and wrote the paper. competing interests there are no competing interests. additional files plots containing the data-model comparison of all 47 green land sites are provided in supplementary file s1 (https://doi.org/10.22008/ fk2/jjq7nh). version 2.0 of the gapslip database is available at zenodo (gowan 2023b; https://doi.org/10.5281/zenodo.8036475) and is updated on github: (https://github.com/evangowan/paleo_sea_level). a spreadsheet with all of the paleo sea-level proxies and indicators for greenland can be found on zenodo (gowan 2023c; https://doi. org/10.5281/zenodo.8036552). a comparison of all the sites in gapslip with various earth and ice-sheet models can be found on zenodo (gowan 2023a; https://doi.org/10.5281/zenodo.7923553). references alley, r.b. et al. 2010: history of the greenland ice sheet: paleoclimatic insights. quaternary science reviews 29, 1728–1756. https://doi. org/10.1016/j.quascirev.2010.02.007 andrews, j.t. & voelker, a.h. 2018: ‘heinrich events’ (& sediments): a history of terminology and recommendations for future usage. quaternary science reviews 187, 31–40. https://doi.org/10.1016/j. quascirev.2018.03.017 baranskaya, a.v., khan, n.s., romanenko, f.a., roy, k., peltier, w. & horton, b.p. 2018: a postglacial relative sea-level database for the russian arctic coast. quaternary science reviews 199, 188–205. 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https://doi.org/10.1371/journal.pone.0259816 https://doi.org/10.1371/journal.pone.0259816 paleo sea-level indicators and proxies from greenland in the gapslip database and comparison with mo 1. introduction 2. sea-level data indicators and proxies 2.1 archives of sea-level data 2.2 data compilation 2.3 vertical interpretation and elevation uncertainties 2.4 age control 3. model-data comparison 3.1 the paleomist reconstruction of greenland 3.2 comparison of calculated sea level with proxies and indicators 4 discussion 4.1 exploring potential solutions to the mismatch 4.2 contributions to global sea level 4.3 improvements to the sea-level indicator and proxy database 5. conclusions acknowledgments additional information references figures fig. 1 map showing the locations of the 47 subregions for which there are data in the gapslip datab fig. 2 paleo sea level and comparison with the reference model at kangerlussuaq subregion. (a) a map fig. 3 basal shear stress values used to reconstruct the greenland ice sheet. fig. 4 difference in ice thickness from the present-day greenland ice sheet in paleomist 1.0, report fig. 5 plots showing sea-level indicators and proxies for selected subregions around greenland, and fig. 6 plots showing sea-level indicators and proxies for selected subregions around greenland and t tables table 1 sea-level proxies and indicators across greenland table 2 reservoir age used to correct marine carbonates geusbulletin monograph surlyk monograph surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 1 of 116 jurassic stratigraphy of east greenland finn surlyk*1 , peter alsen2 , morten bjerager2 , gregers dam2 , michael engkilde3 , carina fabricius hansen4 , michael larsen5 , nanna noe-nygaard1 , stefan piasecki6 , jens therkelsen7 & henrik vosgerau2 1department of geosciences and natural resource management, university of copenhagen, copenhagen, denmark. 2geological survey of denmark and greenland, copenhagen, denmark. 3kgs. lyngby, denmark, 4birkerød, denmark, 5ineos oil and gas, virum, denmark. 6globe institute, section for geobiology, university of copenhagen, denmark. 7moe a/s, vordingborg, denmark abstract the east greenland rift basin comprises a series of jurassic subbasins with different crustal configurations, and somewhat different tectonic histories and styles. the roughly n–s elongated basin is exposed in central and northern east greenland over a length of more than 600 km and a width of up to 250 km. the southernmost exposures are found in the largest subbasin in jameson land, while the northernmost exposures are on store koldewey and in germania land. the focus of the present revision is on the jurassic, but the uppermost triassic and lowermost cretaceous successions are included as they are genetically related to the jurassic succession. the whole succession forms an overall transgressive–regressive megacycle with the highest sea level and maximum transgression in the kimmeridgian. the latest triassic – early jurassic was a time of tectonic quiescence in east greenland. lower jurassic deposits are up to about 950 m thick and are restricted to jameson land and a small down-faulted outlier in southernmost liverpool land. the lower jurassic succession forms an overall stratigraphic layer-cake package that records a shift from rhaetian–sinemurian fluvio-lacustrine to pliensbachian – early bajocian mainly shallow marine sedimentation. onset of rifting in the late bajocian resulted in complete reorganisation of basin configuration and drainage patterns, and the depositional basin expanded far towards the north. post-lower bajocian early-rift deposits are up to about 500–600 m thick and are exposed in jameson land, liverpool land, milne land, traill ø, geographical society ø, hold with hope, clavering ø, wollaston forland, kuhn ø, th. thomsen land, hochstetter forland, store koldewey and germania land. upper jurassic rift-climax strata reach thicknesses of several kilometres and are exposed in the same areas with the exception of liverpool land and germania land. in the southern part of the basin, the upper bajocian – kimmeridgian succession consists of stepwise backstepping units starting with shallow marine sandstones and ending with relatively deep marine mudstones in some places with sandy gravity-flow deposits and injectites. in the jameson land and milne land subbasins, the uppermost jurassic – lowermost cretaceous (volgian–ryazanian) succession consists of forestepping stacked shelf-margin sandstone bodies with associated slope and basinal mudstones and massflow sandstones. north of jameson land, block-faulting and tilting began in the late bajocian and culminated in the middle volgian with formation of *correspondence: finns@ign.ku.dk received: 26 feb 2019 accepted: 20 dec 2019 published: 9 jul 2021 keywords: east greenland, stratigraphy, siliciclastic sediments, uppermost triassic, jurassic, lowermost cretaceous edited by: jon r. ineson (geus, denmark) reviewed by: snorre olaussen (the university centre in svalbard, norway); harald brekke (norwegian petroleum directorate, norway) funding: see page 111 competing interests: none declared additional files: none provided https://orcid.org/0000-0001-6218-9054 https://orcid.org/0000-0003-3180-8857 https://orcid.org/0000-0002-2905-3136 https://orcid.org/0000-0002-7582-5360 surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 2 of 116 www.geusbul let in.org strongly tilted fault blocks, and the succession records continued stepwise deepening. in the wollaston forland – kuhn ø area, the volgian is represented by a thick wedge of deep-water conglomerates and pebbly sandstones passing basinwards into mudstones deposited in fault-attached slope aprons and coalescent submarine fans. the lithostratigraphic scheme established mainly in the 1970s and early 1980s is here revised on the basis of work undertaken over subsequent years. the entire jurassic succession, including the uppermost triassic (rhaetian) and lowermost cretaceous (ryazanian–hauterivian), forms the jameson land supergroup. the supergroup is subdivided into the kap stewart, neill klinter, vardekløft, hall bredning, and wollaston forland groups, which are subdivided into 25 formations and 48 members. many of these are revised, and 3 new formations and 14 new members are introduced. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 3 of 116 www.geusbul let in.org introduction the jurassic of east greenland has a long exploration history and its stratigraphic nomenclature has undergone several phases of revision. many new units have been defined since the naming of the first formations early in the 20th century. the first formal lithostratigraphic schemes covering the whole jurassic succession, including the uppermost triassic and lowermost cretaceous, were established by surlyk et al. (1973) for jameson land, surlyk (1977a, 1978a) for the wollaston forland – kuhn ø area, and callomon & birkelund (1980) and birkelund et al. (1984) for milne land (figs 1–3). these schemes have proved very robust over time, but sedimentological and sequence stratigraphical projects carried out in more recent years have led to improved knowledge and a better understanding of the succession. a new subdivision of the pliensbachian – lower bajocian part of the succession by dam & surlyk (1998) was the first step in the revision and refinement of the older, low-resolution jurassic lithostratigraphy of surlyk et al. (1973). in the present paper, a complete lithostratigraphic scheme is established for the jurassic of east greenland (fig. 3). most of the older units have been revised mainly by elevating the ranks of members and formations following more detailed studies. the stratigraphic scheme presented here includes the uppermost triassic and lowermost cretaceous formations and members because they are genetically related to the jurassic units and because they form the base and top units in jurassic groups, respectively. lithostratigraphic boundaries are normally diachronous and formations are in general rarely genetic sedimentary bodies. formations are thus normally not well suited as a basis for basin analysis and interpretation of the stratigraphic development. this is particularly the case in poorly exposed or subsurface successions. however, in the case of the extensively and excellently exposed jurassic of east greenland the lithostratigraphic units are in most cases close to genetic stratigraphic units. this is because the natural boundaries used in their definition commonly coincide with sequence stratigraphic key surfaces and not only with simple lithological changes. however, the key surfaces have a distinct lithological expression and can be traced laterally over large areas. most of the formations of the older schemes have thus been genetic units from the outset and discrepancy between lithostratigraphy and general sequence stratigraphy is rather minor. it is important to emphasise, however, that although the formational boundaries are commonly sequence stratigraphic key surfaces, it is c c' b' a a' 18°w 16°w 76°n 72°n 22°w 20°w saf pdmf df lle lle cretaceous jurassic triassic permian main faults buried deepseated faults stauning alper fault post-devonian main fault dombjerg fault liverpool land escarpment ittoqqortoormiit/ scoresbysund greenland milne land gåseland traill ø geographical society ø hold with hope clavering ø th.thomsen land kuhn ø hochstetter forland store koldewey germania land kulhøj wollaston forland pd m f sa f d f fig 1 jameson land carl sbe rg fjo rd kong oscar fjord liverpool land hurry inlet scoresbysund 50 km 26°w 22°w24°w 2a 2d 2e 2c 2b 28°w 26°w 24°w volquart boon coast hall bredning 74°n fig. 1 simplified geological map of the east greenland rift basin showing the main faults active during the jurassic, the main areas with outcrops of uppermost triassic – lowermost cretaceous formations, and the location of the detailed geological maps (figs 2a–e). modified from surlyk (2003, fig. 2). http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 4 of 116 www.geusbul let in.org their lithological expression rather than their interpretation that forms the basis for the definition. the area of greenland under focus here is, for simplicity, referred to broadly as east greenland, following most earlier literature on this sedimentary basin complex (see surlyk 2003 and references therein). it should be noted, however, that this geographical area is included in the more rigorously defined regions of central east greenland and north-east greenland as presented by ghisler (1990). a short note on place names is also necessary to avoid confusion due to unclear definitions. jameson land is here understood as the peninsula between scoresby sund and kong oscar fjord, bounded to the west by the north–south-trending valley of schuchert dal and to the east by the klitdal valley, which separates jameson land and liverpool land (figs 1, 2a). the name scoresby land has sometimes been used in the literature for what is here considered north-western jameson land, but scoresby land as originally defined covers a much larger area, including but extending beyond jameson land, and the name is thus not used in this account. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 5 of 116 www.geusbul let in.org geological setting the mesozoic basin of east greenland includes a series of jurassic subbasins with different crustal configurations and somewhat different tectonic histories and styles. the largest of the onshore subbasins, the jameson land subbasin, is located to the south in the area of present-day jameson land and probably continues towards the south below the basalts of the volquart boon coast (fig. 1; larsen & marcussen 1992). the jameson land subbasin is delimited by roughly n–s to nw–se-trending border-faults. the northern boundary is interpreted as being defined by a system of deep-seated nw–se-trending cross-faults down-throwing to the south (surlyk 1978b; dam et al. 1995). the faults probably originated as splays on the main n–s faults bordering devonian oblique-slip pull-apart basins (fig. 1). some major faults have been interpreted to have formed during early triassic rift phases, defining the overall ne–sw-trending triassic basins (guarnieri et al. 2017); seidler et al. (2004) described two early triassic rift phases from nw jameson land controlled by nw– se-trending faults. the latest triassic – earliest cretaceous basin evolution of east greenland was reviewed by surlyk (2003). the uppermost triassic – lower jurassic is restricted to jameson land and southern liverpool land (fig. 2a) and deposition was controlled by thermal contraction under relatively quiet tectonic conditions, accommodated by subsidence along the main basin-margin faults. subsidence became markedly asymmetric in middle–late jurassic times, being greatest towards the west or northwest, but jameson land retained the character of an extensive gently tilted platform undisturbed by significant faulting. a major rift-phase began in the late bajocian and culminated in the middle volgian with the development of tilted fault blocks, especially in the wollaston forland – kuhn ø area (surlyk 2003 and references therein; surlyk & korstgård 2013). jurassic deposition in this area took place in progressively deeper water and water depths in the volgian may have reached almost 1 km along the axes of the hanging-wall basins (surlyk 1978a, 1984). the jurassic succession shows marked onlap onto progressively older strata towards the north and on milne land also towards the west (fig. 2b). this was associated with a younging of its base northwards from jameson land, where an almost complete triassic–jurassic succession is preserved. the traill ø – geographical society ø (fig 2c) area was onlapped in the late bajocian and uppermost triassic – lower middle jurassic sediments are missing. on hold with hope, farther towards the north (fig 2e), lower callovian strata rest directly on the lower triassic wordie creek group (vosgerau et al. 2004a). reworked upper bajocian / lower bathonian ammonites occur in the basal cretaceous pebble lag, however, suggesting that the area had been transgressed earlier in middle jurassic time, but the deposits were subsequently removed by pre-callovian erosion. on milne land, callovian and younger sediments onlap crystalline basement, forming a prominent rocky coast (larsen et al. 2003). on wollaston forland, bathonian deposits overlie upper permian carbonates to the south and crystalline basement to the north (fig. 4). farther north, on kuhn ø, hochstetter forland, store koldewey and germania land (figs 1, 2d, 2e), upper bathonian or lower callovian marine or fluvial sandstones with coal rest on peneplaned crystalline basement. the progressive northwards jurassic onlap has been interpreted to reflect late early jurassic domal uplift in northern east greenland followed by erosion, domal collapse, onset of rifting and transgression in middle jurassic times (surlyk 1977b, 1978b). http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 6 of 116 www.geusbul let in.org 22°w23°w24°w25°w 23°w24°w 72°n 71°30'n 71°n 70°30'n70°30'n 25 km 17 45 4 40 33 35 36 37 32 9 25 27 53 24 13 47 2 6 43 50 54 11 55 30 12 16 5 23 10 42 41 48 46 51 29 18/20 39 52 44 1934 31 1/49 3 28 14 1538 21 227 8 26 olympen fm pelion fm neill klinter gp kap stewart gp triassic charcot bugt fm kap leslie fm basement liverpool land jameson land hurry inlet carls ber g f jord devonian carboniferous permian major sill/dyke a fault ice river quaternary hesteelv fm hartz fjeld fm / pinnadal fm palaeogene basalts palaeogene raukelv fm hareelv fm / bernbjerg fm cretaceous (post-palnatokes bjerg fm) palnatokes bjerg fm lindemans bugt fm 71°nlocalities 1. albuen, 2. antarctic havn 3. astartekløft, 4. coloradodal 5. crinoid bjerg, 6. dusén bjerg 7. elis bjerg, 8. falsterelv 9. fossilbjerget, 10. fynselv 11. gåseelv, 12. goniomyakløft 13. gule horn, 14. hareelv 15. harris fjeld, 16. hesteelv 17. horsedal, 18. innakajik 19. kap hope, 20. kap stewart 21. katedralen, 22. klitdal 23. langelandselv, 24. lepidopteriselv 25. liaselv, 26. lollandselv 27. mikael bjerg, 28. moskusoksekløft 29. muslingeelv, 30. nathorst fjeld 31. neill klinter, 32. olympelven 33. olympen, 34. ostreaelv 35. parnas, 36. pelion 37. pingeldal, 38. primulaelv 39. qupaulakajik, 40. ranunkeldal 41. raukelv, 42. rauk plateau 43. ræveelv, 44. rævekløft 45. rhætelv, 46. salix dal 47. schuchert dal, 48. sjællandselv 49. skævdal, 50. sortehat 51. straight river, 52. tancrediakløft 53. trefjord bjerg, 54. ugleelv 55. zackenberg fig. 2 maps showing the position of place names and localities mentioned in the text. based on the digital greenland geological map at a scale of 1:500 000 (kokfelt et al. 2013). (a): geological and locality map of jameson land. the outcrop area indicated as the pelion formation on this map also includes the fossilbjerget formation. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 7 of 116 www.geusbul let in.org ! ! ½ 25°20'w 70°50'n 20 km 25°40'w localities 56. aldinger elv 57. bay fjelde 58. cardioceraskløft 59. charcot bugt 60. hartz fjeld 61. kosmocerasdal 62. krebsedal 63. kronen 64. mudderbugt 65. parat kløft 66. pernaryggen 67. pinnadal 68. slottet 69. visdal 64 69 57 56 65 63 68 62 67 60 66 58 59 61milne land b 22°w23°w 72°30'n 72°15'n 25 km 71 70 73 74 72 localities 70. bjørnedal 71. mols bjerge 72. steenstrup dal 73. svinhufvud bjerge 74. vælddal traill ø kong oscar fjord c 20 km store koldewey 96 94 97 95 19°w 18°w 75°30'n 75°15'n 76°n localities 94. midter gneisnæs 95. ravn pynt 96. ravn ravine 97. trækpasset d fig. 2 (continued) maps showing the position of place names and localities mentioned in the text. based on the digital greenland geological map at a scale of 1:500 000 (kokfelt et al. 2013). (b) : geological and locality map of south-east milne land. (c): geological and locality map of eastern traill ø. the outcrop area indicated as the pelion formation on this map also includes the bristol elv, fossilbjerget and olympen formations. (d): geological and locality map of store koldewey. the outcrop area indicated as pelion formation on this map also includes the bastians dal, muslingebjerg, payer dal and bernbjerg formations. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 8 of 116 www.geusbul let in.org ½ ! !! ! 21°w 20°w 19°w 75°n 74°30'n 74°n 25 km 74°n 85 90 84 83 81 92 82 80 77 91 78 79 8987 88 93 76 75 86 localities 75. bastian dal 76. bernbjerg 77. cardiocerasdal 78. falkebjerg 79. falskebugt 80. gulelv 81. hohgant 82. jakobsstigen 83. jurakløft 84. kingofjeldet 85. kulhus 86. laugeites ravine 87. niesen 88. palnatoke bjerg 89. payer dal 90. perisphinctes ravine 91. rødryggen 92. stratumbjerg 93. ugpik ravine hochstetter forland wollaston forland hold with hope kuhn ø clavering ø e fig. 2 (continued) maps showing the position of place names and localities mentioned in the text. based on the digital greenland geological map at a scale of 1:500 000 (kokfelt et al. 2013). (e): geological and locality map of the hold with hope, clavering ø, wollaston forland, kuhn ø and hochstetter forland area. the outcrop area indicated as the pelion formation on this map also includes the bastians dal, muslingebjerg, payer dal and jakobsstigen formations. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 9 of 116 www.geusbul let in.org revised lithostratigraphy early comprehensive studies and reviews of the jurassic succession of east greenland were given by maync (1947), donovan (1957), haller (1971), birkelund & perch-nielsen (1976), and surlyk (1990a). a synthesis of the jurassic evolution of east greenland was presented by surlyk (2003). the lithostratigraphic scheme for the jurassic of east greenland was established by surlyk et al. (1973), surlyk (1977a, 1978a), callomon & birkelund (1980) and birkelund et al. (1984). the schemes of surlyk (1977a, 1978a), callomon & birkelund (1980), birkelund et al. (1984) with minor revision, dam & surlyk (1998) and alsgaard et al. (2003) remain essentially unchanged but are included here for the sake of completeness. the other schemes are updated, following sedimentological, biostratigraphic and sequence stratigraphic studies of most of the formations during later years (fig. 3). the area under consideration lies in a remote part of greenland (fig. 1). there is only one permanent settlement (ittoqqortoormiit/scoresbysund; fig. 1) and place names are few, especially in inland areas of large islands and peninsulas. the problem of lithostratigraphic name-giving is thus similar to that in offshore areas where sea birds, explorers, mythological figures or fishing-boat gear have been used in british and norwegian offshore jurassic stratigraphic terminology. in a few cases, names from greek mythology are similarly proposed for new lithostratigraphic units because the nearby mountains pelion, parnas and olympen with greek mythological names have already been used for older units (surlyk et al. 1973; larsen & surlyk 2003). key references are given under the individual units except for units where they are the same as for the formation or group. the place names used in naming the lithostratigraphic units commonly have danish suffixes for river, mountain etc. the most common are elv = river or mountain stream, bjerg = mountain, dal = valley, kløft = ravine, klint(er) = cliff(s). these names as used in the lithostratigraphic scheme should not be translated into english. the spelling of a name of an established stratigraphic unit must not be changed even if the spelling of the place name itself has subsequently been changed. thus, the palnatokes bjerg formation retains its original form (with the genitive ‘s’) despite the official geographical term now being palnatoke bjerg. geographic coordinates are noted for all type sections. many of these were measured at a time when available topographic maps were of poor quality and their precise locations could not be identified with certainty; other localities were positioned in the field on aerial, commonly oblique photographs. in all cases, the positions of the sections are indicated here as precisely as possible. the sedimentary logs were measured by a number of different workers over a relatively long period of time and the style of presentation thus varies depending on the source, in some cases quite strongly. some were measured by sedimentologists and grain sizes are carefully recorded and shown on the horizontal axis. others were measured by non-sedimentologists and although sedimentary thicknesses and positions of samples were accurately recorded, little attention was paid to sedimentary facies and structures. some sections were drawn to show topographic expression or weathering profile along the horizontal axis, rather than grain size, so that cemented beds stand out compared to loose, poorly cemented beds. we have attempted to present all logs in the same style, but in some cases the original log only depicted lithology and grain size by means of ornamentation rather than indicating grain-size variation along the horizontal axis of the log. these logs have been redrawn so that sandstones are labelled ‘sand’ and mudstones as ‘clay’ or ‘silt’ on the x-axis; more precise indications of grain size would have been based on guesswork. a number of new units are defined here. they are described in more detail than those that were formally defined and adequately described in earlier literature (indicated under key references). some of the logs show only part of the unit in question because the base and/ or the top of the unit was not exposed or was inaccessible at the locality; this partial nature of the section is indicated by ‘pars’ on the logs. the ammonite faunal horizons that provide the biostratigraphic framework for many of the units presented in this work refer to the ammonite successions described from milne land (callomon & birkelund, 1980; birkelund et al. 1984) and jameson land (callomon 1993; callomon et al. 2015). note that the numbers of the faunal horizons are given the prefixes mand j-, to clearly differentiate those defined from milne land from those recognised in jameson land (callomon et al. 2015). jameson land supergroup revised unit history. the jameson land group was erected by surlyk et al. (1973). it is here elevated to the rank of supergroup and revised so as to embrace all of the jurassic succession of east greenland, including the genetically related uppermost triassic and lowermost cretaceous strata. type area. jameson land, central east greenland. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 10 of 116 www.geusbul let in.org uppermost triassic – lowermost cretaceous of jameson land group formation member group formation member raukelv fynselv salix dal sjællandselv salix dal sjællandselv hareelv olympen fossilbjerget vardekløft pelion sortehat sortehat ostreaelv ostreaelv neill klinter gule horn gule horn rævekløft rævekløft kap stewart kap stewart jameson land jameson land jameson land supergroup vardekløft hall bredning neill klinter raukelv hesteelvhesteelv hareelv olympen fossilbjerget pelion rhætelv primulaelv innakajik straight river fynselv kokkino prasino langelandselv zeus hades athene goniomyakløft parnas ugleelv trefjord bjerg skævdal nathorst fjeld/harris fjeld/ lepidopteriselv astartekløft horsedal albuen elis bjerg surlyk et al. (1973) this study jurassic – lowermost cretaceous of milne land formation member hennigryggen pinnadal hartz fjeld charcot bugt kap leslie astartedal pernaryggen krebsedal gråkløft cardioceraskløft bays elv aldinger elv kosmocerasdal mudderbugt visdal formation member hennigryggen kronenhartz fjeld pinnadal charcot bugt kap leslie astartedal pernaryggen krebsedal gråkløft cardioceraskløft bays elv aldinger elv kosmocerasdal mudderbugt visdal supergroup group callomon & birkelund (1980) and birkelund et al. (1984) this study rauk plateau muslingeelv crinoid bjerg muslingeelv crinoid bjerg new revised elevated in rank fig. 3 fig. 3 stratigraphical scheme showing older subdivisions together with the revised subdivision presented in this study, covering jameson land, milne land, traill ø, geographical society ø, hold with hope, clavering ø, wollaston forland, kuhn ø, hochstetter forland, store koldewey and germania land. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 11 of 116 www.geusbul let in.org ravn (1911) koch (1929) piasecki et al. (2004) this studysurlyk (1977b) jurassic – lowermost cretaceous of store koldewey and germania land formation/ memberformation formation member supergroup group memberformation jameson land pelion kap hamburg kløft i trækpas vardekløft fm pelion mb kløft i vardekløft bernbjerg spath plateauspath plateau pelion muslingebjerg bastians dal payer dal bernbjerg payer dal jurassic – lowermost cretaceous of hold with hope, clavering ø, wollaston forland, kuhn ø and hochstetter forland surlyk (1977b, 1978b) this study group formation member group formation member wollaston forland jameson land jameson land jameson land wollaston forland palnatokes bjerg lindemans bugt bernbjerg bernbjerg bernbjerg bernbjerg vardekløft vardekløft vardekløft niesen rigi laugeites ravine niesen niesen rigi laugeites ravine jakobsstigen pelion pelion pelion muslingebjerg jakobsstigen pelion muslingebjerg supergroup group formation member group formation member rødryggen albrechts bugt supergroup jameson land wollaston forland wollaston forland vardekløft payer dal bastians dal parnas olympen fossilbjerget bristol elv surlyk (1977b) this study jurassic – lowermost cretaceous of traill ø and geographical society ø lindemans bugt lindemans bugt palnatokes bjerg rødryggen albrechts bugt falske bugt young sund rødryggen albrechts bugt falskebugt young sund ugpik ravine spath plateau midter gneisnæs ravn pynt palnatokes bjerg palnatokes bjerg fig. 3 (continued) stratigraphical scheme showing older subdivisions together with the revised subdivision presented in this study, covering jameson land, milne land, traill ø, geographical society ø, hold with hope, clavering ø, wollaston forland, kuhn ø, hochstetter forland, store koldewey and germania land. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 12 of 116 www.geusbul let in.org thickness. 1100–2400 m in the jameson land subbasin. greater thicknesses are reached in the wollaston forland area. lithology. the kap stewart group at the base is terrestrial and consists of alluvial fan, fluvial and deltaic conglomerates and sandstones, passing basinwards into lacustrine black mudstones. a tidally influenced marginal marine interlude has, however, been recognised in central and northern jameson land. the remainder of the supergroup consists almost exclusively of marine conglomerates, sandstones, heteroliths and mudstones. boundaries. the supergroup unconformably overlies the triassic scoresby land group, which includes the youngest red-beds and dolomites of the jameson land area and older rocks farther north, in places with erosional unconformity. the supergroup is the youngest preserved unit in jameson land. it is overlain by the mudstone-dominated lower cretaceous brorson halvø group in the traill ø, hold with hope and wollaston forland areas. in other areas, the supergroup is unconformably overlain by quaternary deposits or forms the present-day erosional topographic surface. distribution. the supergroup is widely distributed in east greenland (central east greenland and north-east greenland of ghisler 1990), from jameson land and milne land in the south to store koldewey and germania land in the north (fig. 1). chronostratigraphy. rhaetian – lower hauterivian. subdivision. the supergroup is subdivided (from below) into the kap stewart (rhaetian–sinemurian), neill klinter (pliensbachian – lower bajocian), vardekløft (upper bajocian – ryazanian), hall bredning (upper callovian – hauterivian), and wollaston forland (middle volgian – lower hauterivian) groups. kap stewart group revised unit history. the kap stewart formation was erected by rosenkrantz (1929) and revised by surlyk et al. (1973). harris (1937) subdivided the formation into a ‘barren sandstone’ below and a ‘plant-bearing series’ above, and these units to some extent correspond to two of the three new formations of the group. type locality. coastal cliff at kap stewart, south-eastern corner of jameson land, 70°26.7′n, 22°37.8′w. the lower boundary dips below the sea in this area. the upper boundary is well exposed at tancrediakløft, south-eastern jameson land (fig. 2a). for details, see component formations. reference sections. ranunkeldal, horsedal, rhætelv, lepidopteriselv, primulaelv, astartekløft and qupaulakajik (fig. 2a). see under the individual formations. thickness. at outcrop, the group is typically 155–300 m thick, increasing from south to north and from the eastern and western margins to the centre of the basin. in the outlier in south liverpool land, it is only 20 m thick. seismic data show that the thickness may exceed 500 m in the central parts of the basin (dam et al. 1995). lithology. conglomerates, arkosic sandstones, dark plant-bearing mudstones and a few thin coal beds in south-eastern jameson land. towards the north and in central jameson land, thick units of black finely laminated lacustrine mudstones alternate with lake delta sheet sandstones (see figs 5, 7, 10). fossils. abundant plants, spores and pollen. almost barren of invertebrate fossils, but scattered bivalves, gastropods, ostracods, fish scales and spines, coprolites, rare insects and trace fossils occur. depositional environment. mainly terrestrial. depositional settings included alluvial fans, braided and meandering streams, distributary channels and interdistributary bays along the basin margins, and a stratified anoxic lake in the centre of the jameson land subbasin. tidally influenced deposits have, however, been identified in the lower part of the group in central and northern jameson land by clemmensen (1976) and surlyk (unpublished observations at ranunkeldal). boundaries. overlies red marly mudstones and grey carbonates of the norian fleming fjord group. the lower boundary is mainly poorly exposed or obscured by palaeogene volcanic intrusions. the boundary is placed at the base of the lowest yellowish fluvial sandstones or greyish to black mudstones overlying the carbonate deposits of the tait bjerg member or the reddish and greenish mudstones and sandstones of ørsted dal formation, both of the underlying fleming fjord group (clemmensen et al. 2020). along the eastern basin margin, the upper boundary is placed at an erosional unconformity separating polymict pebble conglomerates of the rævekløft formation (neill klinter group) above from sandstones of the kap stewart group below. in the central and northern parts of the basin, the kap stewart group is conformably overlain by bioturbated, fossiliferous sandstones of the elis bjerg member (gule horn http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 13 of 116 www.geusbul let in.org formation, neill klinter group). the boundary is relatively sharp, but no hiatus appears to be present and the sharp transition probably reflects marine inundation of the basin. distribution. jameson land. a small down-faulted outlier occurs in southernmost liverpool land (fig. 1). chronostratigraphy. rhaetian–sinemurian, based on macroflora, spores and pollen. in south-eastern jameson fig. 4 chronostratigraphic scheme of the uppermost triassic – lowermost cretaceous successions of east greenland, showing the main lithologies, depositional environments and tectonostratigraphic sequences. al: albrechts bugt member. as: astartekløft member. at: athene member. b: bernbjerg formation. ba: bastians dal formation. f: falskebugt member. fo: fossilbjerget formation. gu: gule horn formation. ha: hades member. har: hareelv formation. he: hesteelv formation. ho: horsedal member. i: innakajik formation. j: jakobsstigen formation. l: laugeites ravine member. li: lindemans bugt formation. l-n-h: lepidopteriselv, nathorst fjeld and harris fjeld members. n: niesen member. mu: muslingebjerg formation. o: ostreaelv formation. pa: palnatokes bjerg formation. pay: payer dal formation. pe: pelion formation. pr: primulaelv formation. r: rødryggen member. ra: raukelv formation. rh: rhætelv formation. ri: rigi member. ræ: rævekløft formation. sk-t: skævdal and trefjord bjerg members. so: sortehat formation. str: stratumbjerg formation. ug: ugpik ravine member. z: zeus member. ma 125 135 130 145 140 155 150 165 160 175 170 185 180 190 195 200 205 210 ju ra ss ic tr ias sic lo w er u pp er (p ar s) lo w er (p ar s) m id dl e u pp er c re ta ce ou s l l l m he ra str ? r n ? b pay pe ug ba onlaps crystalline basement onlaps upper permian mu j f alpa li l ri har ha at pe+fo so fo sk-t l-n-h o asho gu pr i i prrh ræ l l l l l l m m m l l u u u m u u u u u u l u u barremian hauterivian valanginian berriasian tithonian vo lgi an ry az . sy nrif t m eg as eq ue nc e sy nrif t m eg as eq ue nc e pr erif t m eg as eq ue nc e kimmeridgian oxfordian callovian bathonian bajocian aalenian toarcian pliensbachian sinemurian hettangian rhaetian fig 4 u u l l u no data no data no data jameson landchronostratigraphy s n w e s n wollaston forland – kuhn ø z source rock estuarine/lagoonal sandstones, heteroliths, mudstones and coal beds hiatus/condensed lacustrine deltas, sand-dominated with thin coal beds lacustrine mudstones alluvial/delta plain – paralic, sand-dominated fluvial and estuarine sandstones, conglomerates shallow marine sandstones offshore/basinal mudstones, heteroliths deep marine sandstones deep marine conglomerates coal red marine mudstones http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 14 of 116 www.geusbul let in.org land, along the basin margin, the boundary to the overlying neill klinter group conceals a hiatus corresponding to the sinemurian stage as shown by combined evidence from ammonites, macroflora and palynology. subdivision. the group is subdivided into the innakajik, primulaelv and rhætelv formations. key references. harris (1937 and references therein), rosenkrantz (1942), surlyk et al. (1971, 1973), sykes (1974a), clemmensen (1976), pedersen & lund (1980), dam & christiansen (1990), dam & surlyk (1992, 1993, 1995, 1998), dam et al. (1995), mathiesen et al. (1995), krabbe (1996), lindgreen & surlyk (2000), hesselbo et al. (2002), surlyk (2003), mcelwain et al. (2007, 2009), mcelwain & punyasena (2007), popa & mcelwain (2009), 4 m 3 2 1 0 4 m 3 2 1 0 40 41 42 43 44 45 46 47 48 49 50 m 50 51 52 conglomerate 53 54 55 56 57 58 59 60 m 340° 20° 20° sandstone mudstone (laminated/structureless) compressed logs and branches fig 5 fig. 5 lateral profile of the type section of the innakajik formation, part of a 200 m long measured profile along the coastal cliff at kap stewart (innakajik), jameson land (figs 1, 2a). http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 15 of 116 www.geusbul let in.org mander et al. (2010, 2013), milàn et al. (2012), steinthorsdottir et al. (2012, 2015), kürschner et al. (2014), pott (2014), williford et al. (2014), kelly et al. (2015), hansen et al. (2016), decou et al. (2017). innakajik formation new formation history. the formation corresponds roughly to the ‘barren sandstones’ of harris (1937), see also surlyk et al. (1971) and pedersen & lund (1980). the term ‘barren’ refers to the lack of well-preserved plant fossils. name. after innakajik, the greenlandic name for kap stewart. type section. the coastal cliff at innakajik (kap stewart), 70°26.7′n, 22°37.8′w (figs 2a, 5). neither the lower nor the upper boundary are exposed in this section. reference section. mouth of the ranunkeldal valley (fig. 2a). thickness. 85 m are exposed at the type section, and more than 150 m at ranunkeldal. lithology. disorganised pebble and cobble conglomerates, with common large branches, twigs and organic debris. bedding is extremely irregular and bed boundaries are strongly curved, forming both erosional concave-up or depositional convex-up surfaces. the conglomerates alternate with multistorey pebbly sandstones, thin coaly horizons and localised erosional remnants of dark grey mudstones (figs 5, 6). fossils. rare plants, spores and pollen. depositional environment. alluvial fan and braided to low-sinuosity rivers. boundaries. the lower boundary is rarely exposed. in the area just north of the type section, it appears to be conformable and is placed at the base of yellow sandstones and dark mudstones, overlying reddish and greenish mudstones and sandstones of the ørsted dal formation or the undifferentiated fleming fjord group. farther north, the lower boundary is placed at the base of the lowest yellowish fluvial sandstones or greyish to black mudstones, overlying carbonate deposits of the tait bjerg member of the fleming fjord group. the upper boundary is placed at the base of the lowest sandy fining-upward sandstone–mudstone units or plant-rich mudstones of the primulaelv formation. distribution. south-eastern and western jameson land (figs 1, 2a); the formation wedges out towards the basin centre, interdigitating with the rhætelv formation (fig. 4). chronostratigraphy. rhaetian, based on spores and pollen. fig 6 fig. 6 conglomerates, pebbly sandstones and compressed coalified branches and logs deposited in alluvial fan and braided river environments. black horizontal log (arrowed) is about 1 m long. type section of the innakajik formation, kap stewart, southernmost jameson land (figs 1, 2a). http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 16 of 116 www.geusbul let in.org 0 cl si sand f m c 5 10 15 20 m m 25 30 35 40 45 50 55 60 cl si sand f m c cl si sand f m c cl si sand f m c cl fig 7 si sand f m c 65 70 75 80 70 85 90 95 100 105 n ei ll kl in te r g p pr im ul ae lv fm (p ar s) m pr im ul ae lv fm (p ar s) fig. 7 type section of the primulaelv formation. tancrediakløft, jameson land (figs 1, 2a). the accompanying legend (following page) is applicable to all sedimentary logs in the bulletin. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 17 of 116 www.geusbul let in.org 0 cl si sand f m c 5 10 15 20 m m 25 30 35 40 45 50 55 60 cl si sand f m c cl si sand f m c cl si sand f m c cl fig 7 si sand f m c 65 70 75 80 70 85 90 95 100 105 n ei ll kl in te r g p pr im ul ae lv fm (p ar s) m pr im ul ae lv fm (p ar s) fig. 7 (continued) type section of the primulaelv formation. tancrediakløft, jameson land (figs 1, 2a). the accompanying legend (following page) is applicable to all sedimentary logs in the bulletin. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 18 of 116 www.geusbul let in.org sharp/erosional or irregular sharp planar gradational lithology legend bed contacts random fabric imbricated rootlets plant fragments logs and branches drifted plant stems/logs bivalves gastropods ammonites belemnites brachiopods crinoids echinoderms corals fish shell fragments conglomerate fabric biota arenicolites isp. chondrites isp. curvolithos multiplex diplocraterion isp. diplocraterion parallelum gyrochorte comosa gyrochorte isp. helminthoida isp. helminthopsis magna monocraterion isp. monocraterion tentaculatum nereites isp. ophiomorpha nodosa palaeophycus herberti phoebichnus trochoides phycodes isp. planolites isp. skolithos isp. spirophycus isp. taenidium serpentinum teichichnus isp. thalassinoides isp. tisoa habichi zoophycos isp. escape burrow unidentified sinuous horizontal burrow trace fossils foreset orientation (crosslamination, cross-bedding) crestline orientation of wave ripples bidirectional current indicators current rose cone-in-cone structures ooids palaeocurrents, miscellaneous parallel lamination planar cross-bedding trough cross-bedding cross-bedding with double mud-drapes cross-bedding with pebbles along foresets structureless slump folds rhythmic lamination lenticular bedding wavy bedding flaser bedding sedimentary structures heterolithic sediments conglomerate, sandstone matrix concretion concretionary horizon mudstone clasts carbonaceous material pyritic mudstone glauconite cross-lamination/formset incipient wave ripple cross-lamination wave ripple cross-lamination coarse-grained ripples clinoform orientation hummocky and swaley cross-stratification weak moderate bioturbation intense mud partings heterolith muddy sandstone sandstone pebbly sandstone sideritic rip-up mudstone clasts/conglomerate coal claystone mudstone mudstone fig 7 (continued) legend to fig. 7 http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 19 of 116 www.geusbul let in.org primulaelv formation new formation history. corresponds broadly to the ‘plant-bearing series’ of harris (1937). the macroplant flora indicates that the ‘barren sandstone’ and the lowest 30 m of the ‘plant-bearing series’ belong to the lepidopteris floral zone (rhaetian). above a c. 5 m transition zone, the remaining 55 m of the ‘plant-bearing series’ belong to the hettangian thaumatopteris floral zone (harris 1937). this floral change occurred synchronously with a light carbon isotope excursion, suggesting a causal link between the loss of taxa and the very earliest eruptive phases of the central atlantic magmatic province (hesselbo et al. 2002). name. after the primulaelv river, south-eastern jameson land (fig. 2a). type section. tancrediakløft, 70°30.7′n, 22°37.4′w (figs 2a, 7). reference sections. south side of the astartekløft valley (figs 2a, 8). southern bank at the mouth of the primulaelv river, and ggu slim-hole well core, hareelv no. 303128 (accessible for study at the geological survey of denmark and greenland, copenhagen, denmark). thickness. about 105 m. lithology. pebbly to fine-grained sandstones, which commonly form fining-upwards units, passing up into black to dark grey mudstones, thin coal beds and rootlet horizons. fossils. abundant plants, spores and pollen, rare insects and fish scales. fig 8fig. 8 the primulaelv formation at astartekløft, southern jameson land (figs 1, 2a). the lower part shows alternating, sandstones and dark mudstones deposited in meandering rivers and backswamps. this is overlain by coarse-grained pebbly sandstones and the boundary corresponds to the boundary between the rhaetian lepidopteris zone and the lower jurassic thaumatopteris zone. person at boundary for scale (encircled). http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 20 of 116 www.geusbul let in.org depositional environment. meandering rivers, crevasse splays, delta plain. tidally influenced marginal marine deposits have been identified in central and northern jameson land by clemmensen (1976) and surlyk (unpublished observations at ranunkeldal). boundaries. the lower boundary is defined by the base of the lowest plant-bearing beds, in situ coal beds or well-developed fining-upwards sandstone–mudstone units, overlying the conglomerate-dominated innakajik formation. towards the basin centre, the primulaelv formation interfingers with the rhætelv formation and here the boundary is drawn somewhat arbitrarily at the base of the lowest black organic-rich mudstones. along the eastern basin margin, the upper boundary is defined at the erosional unconformity that marks the base of marine pebbly sandstones and heteroliths of the rævekløft formation (neill klinter group; fig. 9). in the basin centre, the boundary is defined by the base of the lowest bed with marine trace fossils referred to the basal gule horn formation (neill klinter group). distribution. south-eastern jameson land (figs 1, 2a). wedges out towards the basin centre where it interfingers with the rhætelv formation (fig. 4). chronostratigraphy. rhaetian–hettangian, based on macroplants, spores and pollen. probably includes the sinemurian in the basin centre and in the subsurface. rhætelv formation new formation name. after the rhætelv river in northern jameson land (fig. 2a). type section. cliffs north-west of rhætelv, 71°38.8′n, 23°17.2′w (fig. 10). reference sections. lepidopteriselv, horsedal and mouth of ranunkeldal (fig. 2a). thickness. 150–300 m. lithology. black, finely laminated mudstones alternating with sheet sandstones, which are commonly capped by rootlet horizons and thin coaly beds (fig. 11). fossils. fossils are mainly lacking, but rare bivalves, gastropods, ostracods, fish scales and spines, coprolites, plants, spores, pollen and trace fossils do occur. depositional environment. the mudstones were deposited in a deep lake with anoxic bottom conditions, whereas the sheet sandstones represent shallow-water lacustrine deltas. boundaries. in the central and northern parts of the basin, the lower boundary is mainly poorly exposed or obscured by palaeogene volcanic intrusions. the boundary is placed at the base of the lowest yellowish fluvial sandstone or greyish to black mudstones, overlying the carbonate deposits of the tait bjerg member (fleming fig. 9 the boundary between the kap stewart and neill klinter groups marked by an erosion surface with a truncated diplocraterion burrow, overlain by a fossiliferous lag pebble conglomerate of the rævekløft formation, neill klinter group (fig. 7). pencil for scale, 15 cm long. tancrediakløft, jameson land (figs 1, 2a). http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 21 of 116 www.geusbul let in.org fjord group). in the western part of the basin, the lower boundary is poorly exposed but is placed at the base of the lowest coal bed, plant-bearing bed or black mudstone of the rhætelv formation. towards the eastern basin margin, the formation interfingers with the primulaelv formation and the lower boundary is here placed at the base of the lowest fining-upwards sandstone, overlying black organic-rich mudstones. distribution. central and northern jameson land (figs 1, 2a). the rhætelv formation interdigitates with the innakajik and primulaelv formations towards the basin margins (figs 1, 4). chronostratigraphy. rhaetian–sinemurian, based on macroflora, spores and pollen. 0 50 100 130 130 m cl si s pb cl si s pb 4 m 4 m55 m 40 m 45 m rh æt el v fm 200 250 m fig 10 fig. 10 type section of the rhætelv formation, rhætelv, jameson land (figs 1, 2a). modified from dam & surlyk (1993, fig. 3). for legend, see fig. 7; the crossed bar symbol adjacent to the log indicates the position of intrusive sills cutting the section. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 22 of 116 www.geusbul let in.org neill klinter group history. erected as a formation by rosenkrantz (1929) and revised by surlyket al. (1973). elevated to group status by dam & surlyk (1998), who included the sortehat member (elevated to formation status) in the new group. the former sortehat member previously constituted the basal member of the overlying vardekløft formation of surlyk et al. (1973). type locality. the coastal cliffs of neill klinter, south-east jameson land, 70°37.8′n, 22°39.7′w (figs 1, 2a, 12). fig. 11 the rhætelv formation, showing alternating deltaic sheet sandstones and lacustrine deep-water black mudstones. rhætelv, central jameson land (figs 1, 2a). fig. 12 aerial photograph of the neill klinter group at harris fjeld, neill klinter, se jameson land, viewed towards the west (figs 1, 2a). ks: kap stewart group. rk: rævekløft formation. gh: gule horn formation (eb: elis bjerg member, al: albuen member). os: ostreaelv formation (as: astartekløft member, hf: harris fjeld member, sk: skævdal member). the section from the top of the kap stewart group to the top of the harris fjord member is about 160 m thick (see fig. 14). photographed by torben olsen. fig 12 s n ks rk gh os eb al as 100 m hf sk gh os eb al as hf sk http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 23 of 116 www.geusbul let in.org reference sections. harris fjeld, nathorst fjeld, elis bjerg, lepidopteriselv, liaselv, rhætelv, horsedal and ranunkeldal (fig. 2a). see under the individual formations. thickness. 300–450 m, increasing towards the basin centre. lithology. thin conglomerates, cross-bedded sandstones, micaceous heteroliths, sandy mudstones and mudstones. fossils. bivalves, gastropods, ammonites, belemnites, echinoids, crinoids, brachiopods, plants, spores, pollen, dinoflagellate cysts and trace fossils. depositional environment. wave, storm and tidally influenced shallow marine embayment. boundaries. in south-eastern jameson land, the lower boundary is an erosional unconformity, and is placed at the base of the lowest marine conglomerate, cross-bedded sandstone or heterolith, overlying alluvial conglomerates, sandstones or black organic-rich lacustrine mudstones of the kap stewart group. in central and northern jameson land, the boundary is conformable and is placed at the base of the lowest marine sandstone or heterolith with marine trace fossils, overlying black mudstones with sheet sandstones of the rhætelv formation (kap stewart group). the upper boundary is placed at the base of the lowest marine yellow cross-bedded sandstones of the pelion formation (varde kløft group), overlying dark mudstones of the sortehat formation. in south-eastern jameson land, in the coastal cliffs of neill klinter, the upper boundary is in some cases developed as a sharp erosional surface separating black mudstones of the sortehat formation from a thin greyish silty mudstone with ammonites and marine trace fossils rapidly passing upwards into yellow cross-bedded sandstones of the vardekløft group. distribution. restricted to jameson land with a small downfaulted outlier in southern liverpool land (figs 1, 2a). chronostratigraphy. pliensbachian – lower bajocian, based on ammonites, belemnites and dinoflagellate cysts. subdivision. the group is subdivided, from below, into the rævekløft, gule horn, ostreaelv and sortehat formations. key references. rosenkrantz (1929, 1934, 1942), surlyk et al. (1971, 1973), sykes (1974b), dam (1990a, b, 1991), surlyk (1990a, b), doyle (1991), callomon (1993), underhill & partington (1994), krabbe et al. (1994), dam & surlyk (1995, 1998), hansen (1999), koppelhus & dam (2003), koppelhus & hansen (2003), surlyk (2003), nøttvedt et al. (2008), ahokas et al. (2014), kelly et al. (2015), eide et al. (2016), decou et al. (2017). rævekløft formation history. erected as a member by surlyk et al. (1973) and elevated to formation rank by dam & surlyk (1998). type section. rævekløft, south-eastern jameson land, 70°27.4′n, 22°37.9′w (surlyk et al. 1973, fig. 7, section 2; fig. 2a). note that the type section is situated only a few kilometres from the reference section at tancrediakløft (fig. 13) which shows a similar succession in terms of main facies patterns and sedimentological details. the fig 12 s n ks rk gh os eb al as 100 m hf sk gh os eb al as hf sk fig. 12 (continued) aerial photograph of the neill klinter group at harris fjeld, neill klinter, se jameson land, viewed towards the west (figs 1, 2a). ks: kap stewart group. rk: rævekløft formation. gh: gule horn formation (eb: elis bjerg member, al: albuen member). os: ostreaelv formation (as: astartekløft member, hf: harris fjeld member, sk: skævdal member). the section from the top of the kap stewart group to the top of the harris fjord member is about 160 m thick (see fig. 14). photographed by torben olsen. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 24 of 116 www.geusbul let in.org generalised type section log in surlyk et al. (1973, fig. 7, section 2) is thus not reproduced here. reference sections. tancrediakløft (fig. 13), qupaulakajik, harris fjeld, elis bjerg and kap hope (fig. 2a; see dam & surlyk 1998, fig. 7). thickness. 9–20 m. lithology. fossiliferous sandstones, pebbly sandstones and conglomerates. fossils. rich faunas of bivalves, gastropods, ammonites, echinoids and crinoids; trace fossils, spores and pollen. depositional environment. marine shoreface. boundaries. in south-east jameson land, the lower boundary is placed at an erosional unconformity, separating delta-plain sandstones and conglomerates of the kap stewart group below from the overlying marine fossiliferous conglomerates and sandstones of the rævekløft formation. the unconformity marks a hiatus corresponding to the sinemurian stage. the upper boundary is a sharp drowning surface overlain by mudstones and heteroliths of the gule horn formation (elis bjerg member). distribution. south-eastern jameson land (fig. 1). chronostratigraphy. lower pliensbachian, u. jamesoni to lower p. davoei ammonite zones. gule horn formation history. erected as a member by surlyk et al. (1973), elevated to formation rank by dam & surlyk (1998). type section. gule horn, west of carlsberg fjord, 71°20.1′n, 22°43.5′w (surlyk et al. 1973, fig. 8; fig. 2a). the type section illustrated in surlyk et al. (1973) is rather generalised and cannot be redrawn in the same style as the other sections presented here. the reference section at liaselv (dam & surlyk 1998, plate 2) was measured only 5 km sw of the type section on the southern slope of the gule horn mountain and shows the detailed sedimentological features of the formation in the area of the type section. reference sections. neill klinter, harris fjeld (fig. 14), nathorst fjeld, lepidopteriselv, liaselv, ranunkeldal and horsedal (fig. 2a). see under individual members. thickness. 75–185 m, increasing towards the basin centre. lithology. heteroliths and cross-bedded sandstones. fossils. rare bivalves and belemnites, spores and pollen, diverse trace fossils. depositional environment. marine wave and storm-dominated shoreface, tidal channel, subtidal shoal and offshore transition zone. boundaries. along the southeastern margin of the basin, the lower boundary is a drowning surface separating fossiliferous shoreface sandstones and conglomerates of the rævekløft formation below from offshore transition mudstones and heteroliths of the gule horn formation. in the central, northern and western parts of jameson land, the boundary is placed where black lacustrine mudstones of the kap stewart group are overlain by well-sorted fossiliferous siltstones and sandstones of sclsi pb fig 13 g ul e h or n fm (p ar s) ræ ve kl øf t f m n ei ll kl in te r g r ka p st ew ar t g r ( pa rs ) 0 5 10 15 m fig. 13 reference section of the rævekløft formation, tancrediakløft, jameson land (figs 1, 2a). modified from dam & surlyk (1993, fig. 7). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 25 of 116 www.geusbul let in.org the gule horn formation, reflecting a marine inundation of the basin. in the south-eastern part of the basin, along the coastal cliffs of neill klinter, the upper boundary is sharp but non-erosional and placed at the base of storm-dominated shoreface to subtidal sandsheet sandstones of the ostreaelv formation. in the central and northern parts of the basin, the upper boundary is placed at the base of the lagoonal coarsening-upward succession of the ostreaelv formation (horsedal member), overlying cross-laminated and cross-bedded tidal channel and subtidal shoal sandstones of the gule horn formation. distribution. same as the group. chronostratigraphy. not well known. constrained by ages of underlying and overlying formations. dinoflagellate cysts suggest the upper pliensbachian (probably a. margaritatus ammonite zone) to lower toarcian in south-eastern jameson land and ?uppermost sinemurian – upper pliensbachian in northern jameson land. subdivision. the formation is subdivided from below into the elis bjerg and albuen members. elis bjerg member history. erected by dam & surlyk (1998). type section. elis bjerg, north of hurry inlet, 22°40.7′w, 70°56.9′n (dam & surlyk 1998, plate 2, elis bjerg/dusén bjerg section, figs 2a, 15). reference sections. neill klinter, harris fjeld, lepidopteriselv, liaselv, ranunkeldal and horsedal (fig. 2a). thickness. 90–185 m, increasing towards the basin centre. lithology. heteroliths and cross-bedded sandstones. fossils. same as the formation. depositional environment. marine wave and storm-dominated shoreface, tidal channel, subtidal shoal and offshore transition zone. boundaries. the lower boundary is the same as for the formation (see above). along the coastal cliffs of neill klinter, in the south-eastern part of jameson land, the upper boundary is placed at the base of alternating mudstones and well-sorted fine-grained storm-dominated shoreface deposits with massive debris flow deposits of the albuen member. in the northern part of the basin, pb cmf si scl 0 20 40 60 80 100 120 180 200 220 240 260 280 m 140 160 o st re ae lv fm so rt eh at f m (p ar s) g ul e h or n fm n ei ll k lin te r g r h ar ris f je ld m b as ta rt ek lø ft m b el is bj er g m b al bu en m b ræ ve kl øf t f m (p ar s) tr ef jo rd b je rg m b fig 14 fig. 14 reference section of the gule horn formation and type section of the ostreaelv formation, harris fjeld, jameson land (figs 1, 2a, 12). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 26 of 116 www.geusbul let in.org it is placed at the base of the lowest thin coarsening upwards unit of waveand storm-dominated lagoonal sandstone heteroliths rich in rootlet horizons and occasional thin coal beds of the horsedal member. distribution. same as the formation. chronostratigraphy. same as the formation. s si pb cl s si pb cl 0 100 90 80 70 50 m 40 fig 15 30 el is bj er g m b (p ar s) 20 10 m 60 m fig. 15 type section of the elis bjerg member. elis bjerg (dusén bjerg), jameson land (figs 1, 2a). from dam & surlyk (1998, plate 2). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 27 of 116 www.geusbul let in.org albuen member history. erected by dam & surlyk (1998). type section. albuen, hurry inlet, 70°34.1′n, 22°37.8′w, (dam & surlyk 1998, fig. 20; figs 2a, 16). reference sections. qupaulakajik (greenlandic name for albuen), goniomyakløft, astartekløft, moskusoksekløft and harris fjeld (fig. 2a). thickness. from less than 1 m up to 26 m. lithology. alternating mudstones and fine-grained, massive and muddy, commonly pebbly sandstones; sandstones with granite boulders are common. fossils. bivalves, belemnites and trace fossils. depositional environment. marine offshore transition zone influenced by storms and the influx of debris flows. boundaries. the lower boundary is placed where cross bedded sandstones and heterolithic deposits of the uppermost elis bjerg member are overlain by muddy deposits at a sharp, non-erosional boundary. the upper boundary is placed at the base of the lowest cross-bedded tidal channel and subtidal sand-sheet deposits of the astartekløft member. the boundary is sharp and in places erosional along the cliffs of neill klinter and farther northwards at nathorst fjeld and elis bjerg. distribution. south-eastern jameson land. wedges out towards the north. chronostratigraphy. upper pliensbachian, based on palynomorphs. ostreaelv formation history. erected as a member by surlyk et al. (1973), revised and elevated to formation rank by dam & surlyk (1998). type section. the original type section in the ostreaelv valley (surlyk et al. 1973, fig. 12) was poorly chosen; the formation is best represented by the excellent exposures at harris fjeld, southern jameson land, 70°43.4′n 22°41.0′w, which is here designated as a new type section, replacing that at ostreaelv (figs 2a, 12, 14). reference sections. qupaulakajik (greenlandic name for albuen), harris fjeld, nathorst fjeld, lepidopteriselv and horsedal (fig. 2a). see under individual members. thickness. 112–155 m. cl 0 10 20 fig 16 m m qz si pb s al bu en m b (p ar s) fig. 16 type section of the albuen member, albuen, jameson land (figs 1, 2a). from dam & surlyk (1998, fig. 20). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 28 of 116 www.geusbul let in.org lithology. highly variable, dominated by fine to very coarse-grained, commonly cross-bedded sandstones with mud-draped foresets. body fossils common in contrast to the underlying gule horn formation. fossils. bivalves, ammonites, belemnites, brachiopods, crinoids, fish scales and trace fossils. depositional environment. marine. tidal channel, lagoon, terminal lobe, subtidal shoal, shoreface and offshore transition zone. 0 cl si s pb 23 cl si s pb 10 20 m h or se da l m b (p ar s) 30 40 m fig 17 fig. 17 type section of the horsedal member, horsedal, jameson land (figs 1, 2a). from dam & surlyk (1998, fig. 24(i), section d). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 29 of 116 www.geusbul let in.org boundaries. along the coastal cliffs of neill klinter, the lower boundary is sharp and placed at the base of tidal channel and subtidal shoal cross-bedded sandstones of the astartekløft member, overlying alternating shoreface sandstones and mudstones of the albuen member. in the northern and central parts of the basin the lower boundary is placed at the base of wave ripple cross-laminated, parallel-laminated and hummocky cross-stratified sandstones arranged in coarsening-upward units of the horsedal member, overlying cross-laminated and cross-bedded sandstones of the elis bjerg member. the upper boundary is placed at the base of the lowest dark mudstone of the sortehat formation, overlying sandstones of the trefjord bjerg member. it is flat, very sharp and in places pebble strewn, and is interpreted as a coalesced sequence boundary and transgressive surface of erosion. distribution. same as the group. chronostratigraphy. uppermost pliensbachian – lowermost aalenian based on ammonites, belemnites and dinoflagellate cysts. subdivision. the formation is subdivided into the horsedal, astartekløft, lepidopteriselv, nathorst fjeld, harris fjeld, skævdal and trefjord bjerg members. horsedal member history. erected by dam & surlyk (1998). type section. horsedal, 71°52.2′n, 23°14.6′w (dam & surlyk 1998, fig. 24, section d; figs 2a, 17). reference section. cliffs north-west of rhætelv, north-central jameson land (figs 1, 2a). thickness. 50–58 m. lithology. thin coarsening-upward successions of sandstreaked mudstone passing into medium-grained sandstone, commonly capped by rootlet horizons and coaly beds. fossils. plants, fish scales and marine trace fossils. depositional environment. marine. wave-dominated beaches and bayhead deltas prograding into a laterally extensive lagoon. boundaries. in the northern part of jameson land, the lower boundary of the horsedal member is placed at the base of the lowermost thin coarsening-upwards unit of sandy heterolithic waveand storm-dominated lagoonal deposits rich in rootlet horizons and occasional thin coals, overlying cross-bedded sandstones of the elis bjerg member. the upper boundary is erosional and placed at the base of cross-bedded or hummocky cross-stratified and bioturbated shoreface sandstones of the lepidopteriselv member, representing a marine drowning of the horsedal member lagoon. distribution. northern jameson land (figs 1, 2a). chronostratigraphy. probably upper pliensbachian, constrained by the ages of underlying and overlying members. astartekløft member history. erected by dam & surlyk (1998). type section. astartekløft, hurry inlet, 70°36.9′n, 22°40.4′w (dam & surlyk 1998, fig. 27; figs 2a, 18). reference sections. nathorst fjeld, harris fjeld, moskusoksekløft, goniomyakløft, albuen and qupaulakajik (fig. 2a). thickness. 18–43 m. lithology. cross-laminated and cross-bedded sandstones, alternating with thin mudstones. fossils. belemnites, oysters, crinoids and trace fossils. depositional environment. marine. tidal channels and subtidal shoals. boundaries. the lower boundary is placed at the base of the lowest cross-bedded tidal channel and sub-tidal sheet sandstones of the astartekøft member, overlying the muddy shoreface deposits of the albuen member. the boundary is sharp and in places erosional in the south-eastern part of the basin. the upper boundary is a sharp marine drowning surface and is placed at the base of the lowest silty mudstones of the correlative nathorst fjeld and harris fjeld members. at nathorst fjeld, a thin extraformational lag conglomerate rests on the basal erosion surface, interpereted as a marine transgressive surface of erosion. distribution. south-eastern jameson land (figs 1, 2a). chronostratigraphy. dinoflagellate cysts from the base of the member suggest the lower toarcian. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 30 of 116 www.geusbul let in.org lepidopteriselv member history. erected by dam & surlyk (1998). type section. lepidopteriselv on the south slope of trefjord bjerg, carlsberg fjord, 71°15.4′n, 22°36.4′w (dam & surlyk 1998, fig. 30; figs 2a, 19). reference sections. liaselv, horsedal and ranunkeldal (fig. 2a). thickness. 48–60 m. lithology. cross-bedded and hummocky cross-stratified sandstones. fossils. bivalves, belemnites, ammonites, crinoids, dinoflagellate cysts and trace fossils. depositional environment. marine. tidal channels and wave and storm-dominated shoreface. boundaries. at the type locality and at horsedal, sandstones of the lepidopteriselv member rest with a sharp, erosional lower boundary on heterolithic deposits of the astartekløft and horsedal member, respectively. the upper boundary is sharp and placed at the base of muddy bioturbated sandstones of the skævdal member, overlying sandstones of the lepidopteriselv member. distribution. northern jameson land (figs 1, 2a). chronostratigraphy. lower toarcian, d. tenuicostatum ammonite zone, d. semicelatum subzone, based on ammonites. 0 cl si s pb 10 17 m 20 17 30 36 m as ta rt ek lø ft m bas ta rt ek lø ft m b al bu en m b (p ar s) n at ho rs t fj el d m b (p ar s) fig 18 fig. 18 type section of the astartekløft member, astartekløft, jameson land (figs 1, 2a). from dam & surlyk (1998, fig. 27). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 31 of 116 www.geusbul let in.org 0 cl si s pb cl si s pb 10 20 27 m fig 19 30 28 40 50 m le pi do pt er ise lv m b (p ar s) le pi do pt er ise lv m b (p ar s) fig. 19 type section of the lepidopteris elv member, lepidopteriselv, jameson land (figs 1, 2a). from dam & surlyk (1998, fig. 30). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 32 of 116 www.geusbul let in.org nathorst fjeld member history. erected by dam & surlyk (1998). type section. nathorst fjeld, 70°49.1′n, 22°39.5′w (dam & surlyk 1998, fig. 32; figs 2a, 20). reference sections. moskusoksekløft, astartekløft and albuen (fig. 2a). thickness. 29–37 m. lithology. a coarsening-upward succession of alternat ing silty mudstones and laminae/beds of fine-grained wave-ripple cross-laminated and hummocky crossstra tified sand stone, sharply overlain by cross-bedded, 0 cl si s pb fig 20 23 cl si s pb 10 20 23 30 40 m as ta rt ek lø ft m b (p ar s) n at ho rs t f je ld m bn at ho rs t f je ld m b sk æv da l m b (p ar s) fig. 20 type section of the nathorst fjeld member, nathorst fjeld, jameson land (figs 1, 2a). from dam & surlyk (1998, fig. 32). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 33 of 116 www.geusbul let in.org bioturbated fossili ferous coarse-grained sandstones and mud stone-clast conglomerates. fossils. bivalves, ammonites, belemnites, brachiopods, crinoids, vertebrates and trace fossils. depositional environment. marine offshore transition zone to shoreface. boundaries. the lower boundary to the sandstones of the astartekløft member is placed at the base of silty offshore mudstones of the nathorst fjeld member at a sharp marine drowning surface. at nathorst fjeld, a thin extraformational lag conglomerate rests on the boundary, interpreted as a marine transgressive surface of erosion. the upper boundary is placed at the base of the lowest strongly bioturbated shelf sandstones of the skævdal member, overlying shoreface sandstones of the nathorst fjeld member at a drowning surface. distribution. south-eastern jameson land (figs 1, 2a). chronostratigraphy. the belemnite parapassolotheuthis polita suggests the lower toarcian, uppermost h. fal­ ciferum or lowermost h. bifrons ammonite zone (d. com­ mune subzone), while the ammonite dactylioceras sp. suggests the lower toarcian, d. tenuicostatum ammonite zone, d. semicelatum subzone (see discussion in dam & surlyk 1998). palynomorphs suggest the lower toarcian. harris fjeld member history. erected by dam & surlyk (1998). type section. harris fjeld, hurry inlet, 70°43.6′n, 22°41.2′w (figs 2a, 21). dam & surlyk (1998) did not specify which of their four sections at harris fjeld was the type section; section iv of dam & surlyk (1998, fig. 34) is herein designated the type section. thickness. 34–40 m. lithology. stacked, tabular low-angle, heterolithic clinoform beds, up to 18 m thick, which pass laterally into cross-bedded sandstones or silty mudstones. fossils. scarce bivalves, belemnites and trace fossils. depositional environment. marine. ebb-tidal deltas. boundaries. the lower boundary is placed at the base of the lowest thin coarsening-upward clinoform bed of the terminal lobe complex of the harris fjeld member, representing a marine drowning surface. the upper boundary is sharp and placed at the base of the lowest mudstones of the overlying skævdal member; this surface is interpreted as a marine transgressive surface of erosion. distribution. harris fjeld area only (figs 2a, 12). fig 21 si s pbcl 0 5 10 15 20 m h ar ris f je ld m b sk æv da l m b (p ar s) as ta rt ek lø ft m b (p ar s) fig. 21 type section of the harris fjeld member, harris fjeld, jameson land (figs 1, 2a). from dam & surlyk (1998, fig. 34, section iv). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 34 of 116 www.geusbul let in.org n at ho rs t f je ld m b (p ar s) sk æv da l m b sk æv da l m b tr ef jo rd b je rg m b so rt eh at f m (p ar s) m 50 40 30 20 m 10 0 cl si s pb cl si s pb fig. 22 type section of the skævdal member, skævdal, albuen, jameson land (figs 1, 2a). from dam & surlyk (1998, fig. 38). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 35 of 116 www.geusbul let in.org chronostratigraphy. the member is considered as time equivalent to the nathorst fjeld and lepidopteriselv members, suggesting the lower toarcian. skævdal member history. erected by dam & surlyk (1998). type section. skævdal at albuen, hurry inlet, 70°35.0′n, 22°38.0′w (figs 2a, 22). reference sections. nathorst fjeld, lepidopteriselv and horsedal (fig. 2a). thickness. 12–40 m. lithology. bioturbated fossiliferous sandy mudstones and muddy fine-grained sandstones, locally cross-bedded, hummocky cross-stratified or wave rippled. fossils. ammonites, belemnites, bivalves, brachiopods, dinoflagellate cysts and trace fossils. 10 0 20 30 40 m tr ef jo rd b je rg m b tr ef jo rd b je rg m b so rt eh at f m (p ar s) sk æv da l m b (p ar s) scl fig 23 si pb scl si pb fig. 23 type section of the trefjord bjerg member, lepidopteriselv, trefjord bjerg, jameson land (figs 1, 2a). from dam & surlyk (1998, fig. 38). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 36 of 116 www.geusbul let in.org depositional environment. marine. offshore transition zone and lower shoreface. boundaries. the lower boundary is sharp, in places erosional, and is placed at the base of the lowermost muddy sandstones and sandy mudstones, overlying sandstones of the nathorst fjeld, harris fjeld and lepidopteriselv members. the boundary is a transgressive surface of erosion. the upper boundary is placed at the base of the lowest tidal channel sandstones of the trefjord bjerg member. it is a sharp erosional unconformity at lepidopteriselv and horsedal and sharp, but non-erosional along the coastal cliff of neill klinter (see fig. 24), and represents a sequence boundary. distribution. same as group. chronostratigraphy. dating is ambiguous and in need of further work as the precise localisation of recovered macrofossils is uncertain. the belemnite ‘parabrachy­ belus’ subaduncatus suggests the uppermost toarcian, d. levesquei ammonite zone, the ammonite dactylioc­ eras sp. indicates the lower toarcian d. tenuicostatum ammon ite zone, d. semicelatum subzone, and palynomorphs suggest the upper toarcian – lower aalenian. trefjord bjerg member history. erected by dam & surlyk (1998). type section. lepidopteriselv, at the south slope of trefjord bjerg, 71°15.2′n, 22°36.8′w (figs 2a, 23). reference sections. albuen (fig. 22), harris fjeld, nathorst fjeld, ostreaelv and horsedal (fig. 2a). thickness. 21–45 m. lithology. cross-bedded sandstones, in places with claydraped foresets, hummocky cross-stratified sandstones, and sandy bioturbated mudstones (fig. 24). fossils. bivalves, belemnites, ammonites and trace fossils. depositional environment. marine. tidal channel, subtidal shoals, and waveand storm-dominated shoreface. boundaries. at ostreaelv, trefjord bjerg and horsedal, the lower boundary is sharp, erosional and is an important sequence boundary. it is placed at the base of the lowest coarse-grained sandstone of the trefjord bjerg member, overlying muddy sandstone of the skævdal fig. 24 boundary between the dark fine-grained skævdal member and the large-scale cross-bedded sandstone of the trefjord bjerg member, both members of the ostreaelv formation. trefjord bjerg, jameson land (figs 1, 2a). persons for scale at the boundary. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 37 of 116 www.geusbul let in.org member (fig. 24). the boundary is draped by a lag of bored sideritic clasts, extra-formational pebbles, logs and belemnites, interpreted as channel lags. along neill klinter, the boundary to the underlying skævdal member is sharp and non-erosional. the upper boundary to the overlying sortehat formation is one of the most conspicuous lithological boundaries in the jurassic succession. it is placed at the base of black mudstones of the sortehat formation, overlying sandstones of the trefjord bjerg member (fig. 25). it is flat, very sharp and in places pebble strewn, and represents a coalesced sequence boundary and transgressive surface of erosion. distribution. same as group. chronostratigraphy. upper toarcian – lowermost aalenian, based on dinoflagellate cysts. sortehat formation history. erected by surlyk et al. (1973) as the lowermost member of the vardekløft formation. elevated to formation rank by surlyk (1990a, b), and excluded from the vardekløft formation by surlyk (1991). included in the neill klinter group by dam & surlyk (1998). fig. 25 boundary between large-scale cross-bedded sandstones of the trefjord bjerg member and dark mudstones of the sortehat formation. person for scale. immediately north of astartekløft, neill klinter, south-east jameson land (figs 1, 2a). f m c m 0 10 20 30 40 50 60 70 o st re ae lv fm tr ef jo rd b je rg m b (p ar s) so rt eh at f m (p ar s) fig 26 cl/si pbs fig. 26 log of a cored section of the sortehat formation at the type locality, sortehat, jameson land (figs 1, 2a). this slimline core (ggu 303143) is accessible for study at the geological survey of denmark and greenland. for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 38 of 116 www.geusbul let in.org type section. sortehat, ugleelv, southern jameson land, 22°50.3′w, 70°53.8′n (fig. 2a). reference sections. sortehat (fig. 26; a slimline core, drilled at the type locality in 1993, is available for study at the geological survey of denmark and greenland), albuen, goniomyakløft, trefjord bjerg, liaselv (fig. 27), pelion, coloradodal, pingeldal (fig. 2a). thickness. 60–120 m. lithology. dark grey to black mudstones with ironstone concretions, passing upwards into silty mudstones with layers of fine-grained sandstone. fossils. bivalves (oysters), saurian vertebrates, macroplants, belemnites in the upper part, trace fossils, spores, pollen, and dinoflagellate cysts. cl/si f m c pb 0 10 20 30 40 50 60 70 75 m 80 75 90 s fig 27 cl/si f m c pb m 100 110 120 130 140 s o st re ae lv fm tr ef jo rd b je rg m b (p ar s) so rt eh at f m so rt eh at f m pe lio n fm (p ar s) fig. 27 reference section of the sortehat formation, liaselv, jameson land (figs 1, 2a). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 39 of 116 www.geusbul let in.org depositional environment. marine offshore, lower part poorly oxygenated and deposited under brackish marine conditions. boundaries. the lower boundary between the sandstones of the trefjord bjerg member and the black mudstones of the sortehat formation is flat, very sharp, in places pebble strewn and represents a coalesced sequence boundary and transgressive surface of erosion. the upper boundary is placed at the base of cross-bedded marine sandstones of the pelion formation (vardekløft group), overlying black mudstones of the sortehat formation. in south-eastern jameson land, the upper boundary is in some cases developed as a sharp erosional surface separating black mudstones of the sortehat formation from a thin greyish silty mudstone bed with ammonites and marine trace fossils that rapidly passes upwards into yellow cross-bedded sandstones, both facies belonging to the pelion formation. distribution. same as the group. chronostratigraphy. aalenian – lower bajocian, based on dinoflagellate cysts. vardekløft group revised unit history. introduced as a formation by rosenkrantz (1929); revised, described and subdivided into the sortehat, pelion and fossilbjerget members by surlyk et al. (1973), who also gave a historical account. the sortehat member was excluded from the formation by surlyk (1991) and included as a formation in the neill klinter group by dam & surlyk (1998). the vardekløft formation is here elevated to group rank and revised to include the bristol elv, bastians dal, muslingebjerg, pelion, fossilbjerget, payer dal, jakobsstigen, olympen, hareelv, bernbjerg, raukelv and hesteelv formations. type area. jameson land. thickness. the group forms the top unit in most areas of its distribution, so the maximum original thickness is not well known, but probably up to about 650 m. lithology. conglomerates and pebbly sandstones with subordinate coals dominate the bastians dal and bristol elv formations, coals characterise the muslingebjerg formation, sandstones dominate the pelion, jakobsstigen, olympen, payer dal and raukelv formations, whereas the fossilbjerget, bernbjerg, hareelv and hesteelv formations predominantly comprise silty mudstones. fossils. ammonites, bivalves, belemnites, brachiopods, crinoids, plant fragments, dinoflagellate cysts and trace fossils. depositional environment. the group comprises fully marine shoreface to offshore deposits in jameson land (pelion, fossilbjerget and olympen formations). farther north, it is fluvial at the base (bristol elv and bastians dal formations), succeeded by coastal swamp, floodplain and shallow marine sediments (muslingebjerg formation) that in turn are overlain by dominantly marine shoreface and offshore shelf deposits (pelion, payer dal, jakobsstigen, olympen, hareelv, bernbjerg, raukelv and hesteelv formations). boundaries. the lower boundary is placed at the base of marine or fluvial sandstones, sharply overlying black mudstones of the sortehat formation in jameson land, continental upper triassic redbeds and sandstones on traill ø and geographical society ø, lower triassic sandstones and mudstones on hold with hope, upper permian carbonates in southern wollaston forland, and crystalline basement rocks in northern wollaston forland, kuhn ø, hochstetter forland, store koldewey and germania land. the group shows marked onlap onto progressively older rocks from south to north. it forms the top stratum in southern jameson land, and is overlain by lower cretaceous mudstones and sandstones on traill ø, geographical society ø, hold with hope, clavering ø, wollaston forland, kuhn ø and hochstetter forland. details on boundaries are given under the component formations. distribution. from the southern tip of jameson land in the south to germania land in the north (fig. 1). chronostratigraphy. bajocian–ryazanian, based on rich ammonite faunas and dinoflagellate cyst assemblages. subdivision. the group is subdivided (from below) into the bristol elv, bastians dal, muslingebjerg, pelion, fossilbjerget, payer dal, jakobsstigen, olympen, hareelv, bernbjerg, raukelv and hesteelv formations. bristol elv formation history. the formation was erected by therkelsen & surlyk (2004). the deposits of the formation were independently recognised and briefly described by stemmerik et al. (1997) and price & whitham (1997) who included them in their pm1 unit. type section. southern svinhufvud bjerge, traill ø, 72°23.1′n, 23°23.9′w (figs 1, 2c, 28, 29). http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 40 of 116 www.geusbul let in.org 30 20 10 0 320 310 300 290 280 270 260 250 240 220 210 200 190 180 170 160 150 140 130 120 110 scl/si cl/sipb c. 19 m poorly exposed c. 40 m poorly exposed c. 30 m dolerite sill fig 28 m br ist ol e lv fm ? tr ias sic (p ar s) pe lio n fm (p ar s) fl em in g fj or d g p s pb m fig. 28 type section of the bristol elv formation, southern svinhufvud bjerge, traill ø (figs 1, 2c). from therkelsen & surlyk (2004, fig. 3). for legend, see fig. 7. 30 20 10 0 320 310 300 290 280 270 260 250 240 220 210 200 190 180 170 160 150 140 130 120 110 scl/si cl/sipb c. 19 m poorly exposed c. 40 m poorly exposed c. 30 m dolerite sill fig 28 m br ist ol e lv fm ? tr ias sic (p ar s) pe lio n fm (p ar s) fl em in g fj or d fm s pb m http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 41 of 116 www.geusbul let in.org reference section. north-eastern svinhufvud bjerge, southern mols bjerge (fig. 2c). thickness. 80–155 m. lithology. cross-bedded pebble conglomerates and pebbly sandstones intercalated with thin, dark grey mudstones and centimetre-thick coal beds with rootlets. fossils. rare plants and palynomorphs. depositional environment. braided river, floodplain, peat swamp and lake. boundaries. the lower boundary is placed at the base of the lowest yellow sandstone or conglomerate overlying red upper triassic mudstones and coarse-grained sandstones. the upper boundary is placed at the base of the lowermost marine sandstone with trace fossils referred to the pelion formation, and represents a marine flooding surface. distribution. svinhufvud bjerge, mols bjerge and probably also vælddal on traill ø (fig. 2c). chronostratigraphy. uncertain, probably ?lower bajocian, constrained by ammonites of the immediately overlying pelion formation, long-ranging palynomorphs and regional considerations. key reference. therkelsen & surlyk (2004). pe do do do be dotr fig 29 fig. 29 photograph of the type section of the bristol elv formation (be), overlain by the pelion formation (pe), downfaulted against triassic redbeds (tr), and intruded by dark palaeogene dolerite sills and dykes (do), southern svinhufvud bjerge, traill ø (figs 1, 2c). the sill at the base of the formation is about 30 m thick. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 42 of 116 www.geusbul let in.org bastians dal formation history. erected by alsgaard et al. (2003). type section. bastian dal, kuhn ø, 74°53.9′n, 20°20.6′w (alsgaard et al. 2003, fig. 11; figs 1, 2e, 30). thickness. up to about 140 m. lithology. fining-upward units of cobble and pebble conglomerate, passing upwards into parallel-laminated and trough cross-bedded micaceous sandstone (fig. 31). high content of carbonaceous material and thin coal beds (<30 cm thick) occur locally. fossils. spores and pollen. depositional environment. braided and low sinuosity rivers. boundaries. overlies crystalline basement rocks, partly as infill of a valley incised into the basement. the upper boundary is placed at the base of the lowest thick coal bed referred to the muslingebjerg formation, or by marine cross-bedded sandstones of the pelion formation, or by marine sandstones of the payer dal formation on hochstetter forland. distribution. central, and locally southern, kuhn ø, store koldewey (figs 1, 2d, 2e). chronostratigraphy. ?bathonian (palynomorphs indicate a general middle jurassic age); the formation is overlain by upper bathonian sandstones of the pelion formation. key reference. alsgaard et al. (2003). pb cocmf si scl 0 10 20 30 40 50 60 70 80 90 100 m ba st ian s d al fm (p ar s) m us lin ge bj er g fm (p ar s) poorly exposed fig 30 fig. 30 type section of the bastians dal formation, east side of bastian dal, kuhn ø (figs 1, 2e). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 43 of 116 www.geusbul let in.org muslingebjerg formation revised unit history. erected as a member by surlyk (1977a). here elevated to formation rank. type section. kulhus, hochstetter forland, 75°11.5′n, 19°59.5′w (surlyk 1977a, fig. 22; figs 2e, 32, 33). reference sections. bastian dal on kuhn ø (fig. 2e), kulhøj in germania land (fig. 1). thickness. up to about 50 m in central kuhn ø (s.d. andrews & h. vosgerau, unpublished data 2018). lithology. coal beds, up to about 5 m thick, alternating with siltstones, horizontally bedded and cross-bedded sandstones; rootlet beds common. fossils. plant fragments, log imprints, thin oyster bed. depositional environment. coastal mires, floodplain, fluvial channels, lagoon and marine shoreface. boundaries. the lower boundary is not exposed in the type section. elsewhere, it is defined at the base of the lowest thick coal bed overlying crystalline basement rocks, or at the base of the lowest metre-thick coal bed overlying fluvial conglomerates and sandstones of the bastians dal formation. on kuhn ø, the upper boundary is placed at the base of marine sandstones with trace fossils of the pelion or payer dal formations, whereas on store koldewey the upper boundary is placed at the base of dark marine mudstones of the niesen member (lindemans bugt formation). distribution. central kuhn ø, south-western hochstetter forland, store koldewey and germania land (figs 1, 2d, 2e). chronostratigraphy. upper(?) bathonian – middle callovian. strata immediately overlying the highest coal bed on hochstetter forland contain dinoflagellate cysts suggesting the upper callovian p. athleta ammonite zone (piasecki & stemmerik 2004). key references. clemmensen & surlyk (1976), surlyk (1977a), petersen et al. (1998, 2002), petersen & vosgerau (1999), alsgaard et al. (2003), bojesen-koefoed et al. (2012). fig 31 fig. 31 fluvial conglomerates and sandstones, bastians dal, formation, bastian dal, kuhn ø (figs 1, 2e). pencil (encircled) for scale. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 44 of 116 www.geusbul let in.org pelion formation revised unit history. erected as the middle member of the vardekløft formation, which originally included the sortehat, pelion and fossilbjerget members (surlyk et al. 1973). here elevated to formation rank. type section. north slope of the mountain of pelion, north-central jameson land, 71°28.3′n, 23°18.0′w (surlyk et al. 1973, fig. 17; figs 1, 2a, 34, 35). reference sections. goniomyakløft (fig. 42), zackenberg, katedralen, mikael bjerg, trefjord bjerg, olympen, all in jameson land, svinhufvud bjerge on traill ø, cardiocerasdal (see figs 45, 46) and hohgant on wollaston forland and payer dal on kuhn ø (figs 1, 2a, 2c, 2e, see fig. 44). thickness. increases from about 10 m in south-eastern jameson land northwards to 310 m at the type section and possibly up to 650 m at antarctic havn although this section has not been studied in detail and may include the lower part of the olympen formation. at least 240 m on traill ø, possibly up to a few hundred metres on wollaston forland and kuhn ø. in most outcrops, it forms the top of the terrain so only minimum thickness estimates can be given. lithology. cross-bedded, planar bedded or hummocky cross-stratified, medium-grained, micaceous sandstones dominate, with subordinate laminated fine-grained sandstones, coarse-grained and pebbly sandstones. fossils. abundant ammonites in jameson land, in addition to belemnites, bivalves, crinoids, brachiopods and trace fossils, notably tisoa habichi. note that the latter taxon, which in the east greenland literature is referred to as diplocraterion habichi, has been recently revised and is now named tisoa habichi (knaust 2019). t. habichi is a highly characteristic trace fossil of this formation but also of other clean shallow marine jurassic sandstones of east greenland, especially the prasino and fynselv members (raukelv formation) and the kronen member (hartz fjeld formation). depositional environment. marine shoreface and shallow shelf. boundaries. in jameson land, the lower boundary is sharp and placed at the base of the lowest marine sandstone, overlying black mudstones of the sortehat formation (for further details, see sortehat formation). south of the mountain of pelion, the floor of an incised valley locally defines the base of the pelion formation; the boundary is erosional, and is placed at the base of grey, cross-bedded, medium-grained sandstones rich in coal intraclasts, overlying dark grey, laminated mudstones of the sortehat formation. north of jameson cl si f m s c pb 0 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 m fig 32 se am 2 se am 3 mu sli ng eb jer g f m (p ar s) fig. 32 type section of the muslingebjerg formation, southern hochstetter forland (figs 1, 2e). from petersen et al. (1998, fig. 2). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 45 of 116 www.geusbul let in.org land, the pelion formation overlies fluvial pebbly sandstones of the bristol elv formation on traill ø, and fluvial conglomerates and sandstones of the bastians dal formation or coal-bearing sandstones of the muslingebjerg formation on kuhn ø. the lower boundary is placed at the base of the lowest yellow marine sandstone overlying lower triassic sandstones and mudstones on hold with hope, upper permian carbonates on southern wollaston forland, and crystalline basement rocks in northern wollaston forland, kuhn ø and store koldewey. the upper boundary is strongly diachronous, younging from south to north. it is sharp and placed at the base of grey silty mudstones of the fossilbjerget formation in jameson land and on traill ø, at the base of heteroliths and coaly mudstones of the jakobsstigen formation on wollaston forland, and beneath sandstones of the payer dal formation on hold with hope, northern wollaston forland and kuhn ø. in some areas on traill ø and store koldewey, it is overlain with erosional unconformity by lower – mid-cretaceous mudstones. distribution. same as the group, with the exception of germania land where the pelion formation is absent. chronostratigraphy. upper bajocian – middle callovian, c. borealis – k. jason ammonite zones (faunal horizons j-1 to j-36 of callomon 1993), based on ammonites. the uppermost strata are poorly fossiliferous and the age relations to the overlying olympen, jakobsstigen, muslingebjerg and payer dal formations are not always certain. at jurakløft, northern wollaston forland, the upper part of the formation contains dinocysts that suggest a late callovian age. subdivision. the formation includes the ugleelv, parnas and spath plateau members. key references. spath (1932), maync (1947), donovan (1957), haller (1971), surlyk et al. (1973), surlyk (1977a, 1990a, b, 1991), surlyk & clemmensen (1983), heinberg & birkelund (1984), callomon (1993, 1994), surlyk et al. (1993), engkilde (1994), engkilde & surlyk (1993, 2003), price & whitham (1997; thickness estimates probably overestimated, see surlyk & noe-nygaard 2001a), surlyk & noe-nygaard (2000), alsen & surlyk (2004), vosgerau et al. (2004a, b), callomon et al. (2015), bjerager et al. (2020). ugleelv member new member history. this new member is defined to describe a variant of the pelion formation in south-east jameson land. the upper part of the member in the type section is dominated by a prominent high-angle clinoform set that has been figured by surlyk et al. (1973, fig. 19), heinberg & birkelund (1984, figs 13–14) and engkilde & surlyk (2003, fig. 27). name. after the ugleelv river in south-eastern jameson land (fig. 2a). type section. katedralen, on the southern side of the ugleelv river valley, 70°53.3′n, 22°52.9′w (figs 2a, 36). reference section. ræveelv (fig. 2a). thickness. 110–170 m. fig. 33 type section of the muslingebjerg formation, showing light-coloured, shallow marine sandstones sandwiched between thick coal beds, southern hochstetter forland (figs 1, 2e). http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 46 of 116 www.geusbul let in.org 0 15 20 25 30 40 35 55 60 cl si pbsand f m c 70 80 65 75 85 90 95 100 105 110 115 120 125 130 140 150 155 165 170 175 180 185 190 195 200 205 210 215 220 225 230 235 240 245 250 255 260 270 275 280 290 330 335 fig 34 338 m 285 295 265 145 135 160 cl si pbsand f m c cl si pbsand f m c cl si pbsand f m c cl si pbsand f m c cl si pbsand f m c pe lio n fm (p ar s) fig. 34 type section of the pelion formation, pelion, north-central jameson land (figs 1, 2a). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 47 of 116 www.geusbul let in.org 0 15 20 25 30 40 35 55 60 cl si pbsand f m c 70 80 65 75 85 90 95 100 105 110 115 120 125 130 140 150 155 165 170 175 180 185 190 195 200 205 210 215 220 225 230 235 240 245 250 255 260 270 275 280 290 330 335 fig 34 338 m 285 295 265 145 135 160 cl si pbsand f m c cl si pbsand f m c cl si pbsand f m c cl si pbsand f m c cl si pbsand f m c pe lio n fm (p ar s) fig. 34 (continued) type section of the pelion formation, pelion, north-central jameson land (figs 1, 2a). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 48 of 116 www.geusbul let in.org lithology. characterised by one (at katedralen) or two (at ræveelv) prominent, cliff-forming, high-angle clinoform-bedded sandstones, which form the upper part of the member (fig. 37). the lower part consists of structureless, cross-laminated, cross-bedded and hummocky cross-stratified very fineto medium-grained sandstones. the middle part of the member in the type section comprises stacked coarsening-upward siltstone–sandstone cycles. fossils. abundant ammonites and belemnites, bivalves, crinoids, brachiopods, rare gastropods, trace fossils and wood fragments. fossils are commonly concentrated in distinct layers. depositional environment. shallow marine shelf and shoreface. upper part represents a tidally influenced shelf-margin wedge. boundaries. in the type section, the lower boundary is placed at the base of light, silty fine-grained structureless to faintly bedded siltstones and sandstones, overlying a dark, silty, fine-grained, laminated sandstone bed of the mudstone-dominated sortehat formation with a sharp, planar contact. the upper boundary is placed where soft reddish-grey, silty, laminated, fossil-rich and bioturbated mudstones of the fossilbjerget formation overlie light grey sandstones of the pelion formation. this boundary commonly weathers out as a marked ammonite and belemnite-strewn, marine drowning surface (fig. 37), regionally reflecting a northward retreat of the shoreface environment. distribution. the area around inner ugleelv and ræveelv, south-eastern jameson land (figs 1, 2a). chronostratigraphy. upper bajocian, c. borealis – basal c. pompeckji ammonite zones (faunal horizons j-1 to j-4 of callomon 1993), based on ammonites. the lower sandstone and middle siltstone–sandstone units both belong to the c. borealis zone; the upper clinoform-bedded unit belongs to the c. indistinctus to c. pompeckji zones. parnas member new member history. first recognised as a regressive sandstone wedge in the upper levels of the previous pelion member fig. 35 type section of the pelion formation. persons for scale. pelion, central jameson land (figs 1, 2a). the vertical sandstone cliff immediately above the geologists is c. 10 m high. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 49 of 116 www.geusbul let in.org by heinberg & birkelund (1984). erected here as a new member of the pelion formation. name. after parnas in north-central jameson land (fig. 2a). type section. north-east slope of the olympen mountain, 71°27.0′n, 23°28.0′w (figs 2a, 38). reference section. parnas (fig. 2a). thickness. about 30 m. lithology. very fineto coarse-grained sandstones forming metreto decimetre-scale coarsening-upward cycles. fossils. abundant ammonites and belemnites, rare bivalves, trace fossils and wood fragments. depositional environment. shallow marine shelf to shoreface. boundaries. the member forms a local prograding sandy wedge in the otherwise silty mudstone-dominated fossilbjerget formation. the lower boundary is placed at the base of a grey, silty to sandy, coarsening-upward unit, about one metre thick, with hummocky cross stratification in the top part and a rich marine fauna of belemnites, ammonites, bivalves, gastropods and brachiopods, 0 20 40 60 80 100 120 140 180 200 160 m pe lio n fm fig 36 fo ss ilb je rg et f m (p ar s) so rt eh at f m (p ar s) u gle el v m b si f m c pbcl s fig. 36 type section of the ugleelv member, ugleelv, jameson land (figs 1, 2a). for legend, see fig. 7. ha fo ug fig. 37 north slope of the ugleelv valley, southern jameson land, a classical middle jurassic ammonite locality (figs 1, 2a). the lowermost unit is the clinoform-bedded, up to 30 m thick, uppermost yellow sandstone of the ugleelv member, pelion formation (ug; about 20 m exposed at this locality). the lowest few metres of the overlying fossilbjerget formation has yielded a super-abundant fauna of well-preserved ammonites of the c. pompeckji zone. the change in slope between the light grey or brown-weathering mudstones and the slightly steeper slope concealing black mudstones and sandstone injectites marks the boundary between the fossilbjerget (fo) and hareelv formations (ha) (boundary indicated). http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 50 of 116 www.geusbul let in.org sharply overlying grey, very fine-grained silty sandstones of the fossilbjerget formation. the boundary records a short-lived departure from the dominant back-stepping, northwards younging boundary between the pelion and fossilbjerget formations to forward-stepping of progradational depositional units in northern jameson land and on traill ø. the upper boundary is placed at the base of grey, evenly laminated to faintly bedded silty sandstones of the upper fossilbjerget formation. it is a major drowning surface that marks the onset of the final back-stepping of the pelion–fossil bjerget couplet in the area. distribution. exposed on the olympen, parnas and pelion mountains and on the northern slopes of fossilbjerget in north-central jameson land, in the bjørnedal valley of south-eastern traill ø (alsen & surlyk 2004), and east of the steenstrup dal valley on southern traill ø (vosgerau et al. 2004b). wedges out towards the south from the area of the type section and is not present at mikael bjerg (figs 2a, 2c). chronostratigraphy. lower callovian, c. apertum – c. norden skjoeldi ammonite zones (faunal horizons j-24 to j-30 of callomon 1993), based on ammonites. spath plateau member history. erected by vosgerau et al. (2004a). type section. composite section where the lower and upper boundaries are exposed in two different ravines on the east side of the gulelv river valley, northern hold with hope, 73°54.95′n 21°12.5′w (lower boundary), 73°54.6′n, 21°12.5′w (upper boundary; vosgerau et al. 2004a, fig. 5; figs 1, 2e, 39, 40). thickness. 155 m. lithology. the member is dominated by coarsening-upward units of sandy heteroliths gradationally or erosionally overlain by cross-bedded fineto coarse-grained sandstones with common mudstone-draped foresets. a silty, very fine-grained sandstone unit (c. 6 m thick), locally with scattered fine pebbles in the lowermost 0.5 m, forms the basal part of the member. fossils. ammonites, belemnites, bivalves, gastropods, dinoflagellate cysts, silicified and coalified wood, leaf fragments. depositional environment. tidally influenced deltas with distributary channels and mouth bars in the lower–middle shoreface. boundaries. on hold with hope, the lower boundary is a sharp drowning surface marked by an abrupt change cl si f m s pb fig 38 c f c 60 50 40 30 20 10 0 m fo ss ilb je rg et f m (p ar s) fo ss ilb je rg et f m pe lio n fm pa rn as m b pe lio n fm (p ar s) fig. 38 type section of the parnas member, olympen, jameson land (figs 1, 2a). the lower boundary of the member is placed where the backstepping motif of the pelion–fossilbjerget formations changes to the forestepping motif of the progradational parnas member. from engkilde & surlyk (2003, fig. 33). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 51 of 116 www.geusbul let in.org from coarse-grained sandstones of the lower part of the pelion formation (informally named ‘lower sandstone unit’ in vosgerau et al. 2004a) to dark brown, silty, finegrained sandstones at the base of the member. this basal fine-grained unit forms a 6 m thick marker bed, situated 30–40 m above the base of the pelion formation. the upper boundary is a sharp erosion surface overlain by pebbly sandstones and sandy heteroliths of the payer dal formation. on store koldewey, it is locally overlain with erosional unconformity by the ravn pynt member of the palnatokes bjerg formation. distribution. the spath plateau member is only recognised in northern hold with hope and on store koldewey (figs 1, 2d, 2e). chronostratigraphy. lower–middle callovian, p. koenigi – pre-p. athleta ammonite zones, based on ammonites and dinoflagellate cysts. sp at h pl at ea u m b sp at h pl at ea u m b ssi pb co cl ssi pb co cl ssi pb co cltr ias sic pe lio n fm pe lio n fm be rn bj er g fm be rn bj er g fm pa ye r d al fm c re ta ce ou s ? fig 39 120 110 100 90 80 70 60 50 40 30 20 10 0 250 240 230 220 210 200 190 180 170 160 150 140 130 360 350 340 330 260 m m m no exposure no exposure no exposure fig. 39 type section of the spath plateau member, gulelv, hold with hope (figs 1, 2e). from vosgerau et al. (2004a). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 52 of 116 www.geusbul let in.org fossilbjerget formation revised unit history. erected as the upper member of the vardekløft formation by surlyk et al. (1973). here elevated to formation rank. type section. fossilbjerget, 71°15.1′n, 23°06.3′w (surlyk et al. 1973, fig. 23; figs 2a, 41). reference sections. goniomyakløft, zackenberg, katedralen, mikael bjerg and pelion in jameson land and bjørnedal on traill ø (figs 1, 2a, 2c, see fig 50). thickness. 80–120 m. lithology. silty, highly micaceous mudstones, with subordinate fine-grained, locally glauconitic, sandstones, which become mediumto coarse-grained northward on traill ø. the sandstones are stacked in coarsening-upwards units. horizons with phosphatic and calcareous concretions are common throughout. fossils. ammonites in profusion, bivalves, belemnites, dinoflagellate cysts, tree trunks and trace fossils. depositional environment. offshore marine. boundaries. in central jameson land, the lower boundary is placed at the base of soft, reddish-grey, silty, laminated, fossil-rich, bioturbated mudstones, overlying massive, light-grey, fine-grained to conglomeratic sandstones of the pelion formation. the boundary commonly weathers out as a marked ammoniteand belemnite-strewn, marine ravinement surface (fig. 37). the lower boundary youngs from south to north and represents a northward back-stepping succession of drowning surfaces. the upper boundary is placed at a sharp contact at the base of laminated, black mudstones of the hareelv formation in southern jameson land (figs 37, 50), and at a sharp contact at the base of light grey, structureless sandstones of the athene member (olympen formation) in central and northern jameson land. on traill ø, the fossilbjerget formation is overlain by sandstones of the zeus member (olympen formation). distribution. mainly known from jameson land but has been traced northwards to traill ø. the sandy parnas member of the northwards backstepping pelion formation forms a southward tapering wedge within the fig 40 lower cretaceous payer dal fm pelion fm (spath plateau mb) pelion fm(lower part) bernbjerg fm lower cretaceous fault n s fig. 40 aerial photograph of the type section of the c. 155 m thick spath plateau member, pelion formation, exposed in a fault block on the eastern side of the gulelv river, northern hold with hope (figs 1, 2e). a major drowning surface separates the member from the basal sandstone unit of the pelion formation. the boundary to the overlying payer dal formation is a sharp erosion surface. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 53 of 116 www.geusbul let in.org upper levels of the fossilbjerget formation in north-central jameson land and on south-eastern traill ø (fig. 1). chronostratigraphy. upper bajocian – lower upper callovian, c. pompeckji – p. athleta zones (faunal horizons j-3 to j-37 of callomon 1993), based on ammonites. boundaries are strongly diachronous, both younging to the north. subdivision: the formation contains a new member, the goniomyakløft member. key references. birkelund et al. (1971), surlyk et al. (1973), heinberg & birkelund (1984), poulsen (1985), callomon (1993), engkilde & surlyk (2003), alsen & surlyk (2004), vosgerau et al. (2004b), callomon et al. (2015). goniomyakløft member new member history. recognised as a sharp-based unit of soft, light grey weathering mudstones, about 20 m thick, forming the uppermost part of the vardekløft formation (of previous usage) in southern jameson land (surlyk et al. 1973); corresponds broadly to the ‘wood beds’ of callomon (1993). type section. goniomyakløft, hurry inlet, 70°35.3′n, 22°38.5′w (surlyk et al. 1973, fig.14, section 6; figs 1, 2a, 42). reference sections. zackenberg and the east slope of the katedralen mountain, southern jameson land (callomon 1993, fig. 4; figs 1, 2a). thickness. about 20 m. lithology. light grey mudstones with large, greenish, irregular, extremely hard concretions commonly with a nucleus formed by a silicified log. fossils. ammonites and wood. depositional environment. offshore marine. boundaries. the lower boundary is placed where light grey to greenish, soft, laminated mudstones abruptly overlie harder, silty, dark-grey, laminated to thinly bedded mudstones with fine-grained sandstone intercalations of the lower fossilbjerget formation. the boundary is sharp and conceals a major hiatus corresponding to faunal horizons j-24 to j-34 of the c. apertum, c. norden­ skjoeldi and lowermost p. koenigi ammonite zones of callomon (1993). the upper boundary is sharp and placed at the base of laminated black mudstones of the hareelv formation. distribution. southern jameson land. chronostratigraphy. top lower callovian – lower upper callovian, s. calloviense – p. athleta zones, based on ammonites. o lym pe n fm (p ar s) fo ss ilb je rg et f m pelion fm 0 cl/si s 50 100 m fig 41 0 5 10 m cl/si s fig. 41 type section of the fossilbjerget formation, fossilbjerget, jameson land (figs 1, 2a). modified from surlyk et al. (1973, fig. 23). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 54 of 116 www.geusbul let in.org payer dal formation history. erected by alsgaard et al. (2003) for the upper part of the pelion member of surlyk (1977a) on kuhn ø. the sections referred here to the payer dal formation in th. thomsen land and on hochstetter forland were previously referred to the pelion and jakobsstigen members by sykes & surlyk (1976) and surlyk (1977a), respectively. the succession on store koldewey was previously referred to as kløft i formation by koch (1929) and this was followed by surlyk (1977a). this formation h ar ee lv fm (p ar s) fo ss ilb je rg et f m g on io m ya kl øf t m b pe lio n fm so rt eh at f m 0 50 100 150 200 250 m ostreaelv fm (pars) n ei ll kl in te r g r va rd ek lø ft g r fig 42 cl/si s fig. 42 type section of the goniomyakløft member, goniomyakløft, jameson land (figs 1, 2a). modified from surlyk et al. (1973, fig. 14). for legend, see fig. 7. 160 150 130 120 110 100 80 70 60 20 m pe lio n fm (p ar s) pa ye r d al fm (p ar s) s f m c sicl pb fig 43 fig. 43 type section of the payer dal formation, western slopes of kingofjeldet on the eastern side of the payer dal valley, kuhn ø (figs 1, 2e). from alsgaard et al. (2003, fig. 17). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 55 of 116 www.geusbul let in.org was abandoned by piasecki et al. (2004) who referred the lower sandstone and upper mudstone units of the kløft i formation of koch (1929) to the payer dal and bernbjerg formations, respectively (figs 1, 2d, 2e). type section. western slopes of kingofjeldet, eastern side of payer dal, kuhn ø, 74°44.9′n, 20°15.3′w (alsgaard et al. 2003, fig. 17; figs 1, 2e, 43, 44). reference sections. ugpik ravine, western side of payer dal, kuhn ø and store koldewey (figs 1, 2d, 2e). thickness. up to 200 m. lithology. sandy heteroliths, fineto coarse-grained, cross-bedded sandstone and pebbly sandstone. a discrete coarsening-upward unit of silty mudstones and sandy heteroliths characterises the lowermost part of the formation at payer dal, kuhn ø. fossils. bivalves, belemnites, dinoflagellate cysts and trace fossils. depositional environment. tidally influenced marine shoreface. boundaries. on wollaston forland and kuhn ø, the lower boundary is placed at the base of siltstones or heteroliths, overlying cross-bedded or structureless, carbonate-cemented sandstones of the pelion formation at a major drowning surface. on hochstetter forland, the boundary is placed at a similar drowning surface, separating the uppermost coal bed of the muslingebjerg formation from fine-grained, silty yellow sandstones of the payer dal formation. the upper boundary is placed where black mudstones and siltstones of the bernbjerg formation overlie yellow sandstones. fig 44 top payer dal fm (projected) payer dal fm pelion fm 10 m se nw fig. 44 field photograph of the type section of the payer dal formation. a sharp drowning surface separates the payer dal formation (c. 200 m thick) from the underlying pelion formation. the boundary with the overlying bernbjerg formation is exposed about 1 km east of this locality and on the western slopes of the payer dal valley, kuhn ø (figs 1, 2e). http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 56 of 116 www.geusbul let in.org distribution. hold with hope, wollaston forland, kuhn ø, th. thomsen land, hochstetter forland and store koldewey (figs1, 2d, 2e). chronostratigraphy. uppermost callovian – middle oxfordian in the wollaston forland – kuhn ø area, based on rare ammonites and dinoflagellate cysts. on store koldewey, upper oxfordian – lower kimmeridgian, a. serratum – r. cymodoce zones, based on ammonites. key references. maync (1947), sykes & surlyk (1976), surlyk (1977a), surlyk & clemmensen (1983), vosgerau (1997), alsgaard et al. (2003), piasecki et al. (2004), vosgerau et al. (2004a). jakobsstigen formation revised unit history. erected as a member of the vardekløft formation by surlyk (1977a) and here elevated to the rank of formation. type section. the northern slope of the cardiocerasdal valley, wollaston forland, named jakobsstigen after a conspicuous flower, 74°27.2′n, 20°14.5′w (surlyk 1977a, fig. 16, section 38; figs 1, 2e, 45–47). reference sections. two previously defined reference sections for the jakobstigen member of previous usage fig 45 s f m c cl/si scl/si scl/si f m c f m c pe lio n fm jak ob ss tig en f m jak ob ss tig en f m be rn bj er g fm u gp ik r av in e m b 180 170 160 150 140 130 120 120 110 100 90 80 70 60 60 50 40 30 20 10 0 m mm fig. 45 type section of the jakobsstigen formation, cardiocerasdal, wollaston forland (figs 1, 2e). from surlyk (1977a, fig. 16). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 57 of 116 www.geusbul let in.org in th. thomsen land (surlyk 1977a, sections 2, 3) are abandoned herein as they are re-assigned to the ugpik ravine member of the bernbjerg formation. thickness. up to 128 m. lithology. strongly bioturbated heteroliths and cross bedded sandstones, forming coarsening-upward units alternating with thin black mudstones. fossils. bivalves, belemnites, rare ammonites and trace fossils. depositional environment. marine shoreface and possibly coastal-plain lakes or interdistributary bays. boundaries. the lower boundary is placed at the base of sandy heteroliths, sharply overlying cross-bedded, coarsegrained sandstones of the pelion formation. the upper boundary is placed at the base of dark mudstones and fine-grained heteroliths of the bernbjerg formation. both boundaries are prominent marine drowning surfaces. distribution. wollaston forland (figs 1, 2e). chronostratigraphy. lower–middle oxfordian based on ammonites. key references. surlyk (1977a), surlyk & clemmensen (1983), bojesen-kofoed et al. (1997), vosgerau et al. (2000). fig 46 bernbjerg fm wsw ene ugpik ravine mb jakobsstigen fm pelion fm ~20 m fig. 46 aerial photograph of the type section of the jakobsstigen formation (arrowed), viewed towards the nw. the formation is underlain by the pelion formation and overlain by the ugpik ravine member of the bernbjerg formation. the succession shows three major drowning surfaces marking backstepping events: (1) the pelion formation – jakobsstigen formation boundary, (2) the jakobstigen formation – ugpik ravine member (bernbjerg formation) boundary, and (3) the boundary between the heterolithic ugpik ravine member and the dark mudstones of the upper part of the bernbjerg formation. north slope of the cardiocerasdal valley, sw wollaston forland (figs 1, 2e). http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 58 of 116 www.geusbul let in.org olympen formation history. erected by surlyk et al. (1973). type section. olympelven, central jameson land, 71°24.4′n, 23°34.8′w (surlyk et al. 1973, fig. 24a; figs 1, 2a, 48). reference sections. summit (altitude 1020 m) 4 km southwest of the olympen mountain top, parnas and mikael bjerg, all in central jameson land, and vælddal, southeast traill ø (figs 1, 2a, 2c). thickness. forms the top unit in the type area, minimum thickness 150 m in the type section, 300 m at parnas, wedges out southward in the hurry inlet area. the formation thins towards the east on traill ø, where the thickness is estimated to be c. 250 m at bjørnedal and c. 100 m at vælddal. lithology. in the exposures in central jameson land, the formation comprises (from below): massive sandstones of the athene member, mudstones with sandy interbeds of the hades member, and massive and cross-bedded sandstones of the zeus member. on traill ø, cross-bedded sandstones and intervening mudstone intervals are referred to the zeus member. fossils. rare ammonites, bivalves, belemnites, dinoflagellate cysts, plant fragments, fossil wood and trace fossils. fig 47 ja pe fig. 47 the jakobstigen formation (ja) on the ridge immediately west of the type section (see fig. 46). yellow sandstones of the pelion formation (pe) at the base. person (encircled) for scale. north slope of the cardiocerasdal valley, sw wollaston forland (figs 1, 2e). 0 50 100 150 m fo ss ilb je rg et f m (p ar s) o lym pe n fm at he ne m b h ad es m b ze us m b (p ar s) mud f m c gr sand fig 48 fig. 48 type section of the olympen formation and its constituent athene, hades and zeus members, olympelven, jameson land (figs 1, 2a). from larsen & surlyk (2003, fig. 2). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 59 of 116 www.geusbul let in.org depositional environment. marine shelf, shelf-margin wedge, slope, base-of-slope and basin. boundaries. in the type section, the base of the formation (a prominent sequence boundary) is placed beneath the lowest massive sandstone bed, overlying silty mudstones of the fossilbjerget formation. the upper boundary is not exposed in the type section but in south-central jameson land, thin-bedded sandstones of the olympen formation are sharply overlain by black mudstones with massive sandstone injectites of the hareelv formation. this represents a marked marine drowning surface. on traill ø, sandstones referred to the olympen formation (zeus member) may be conformably overlain by black mudstones of the bernbjerg formation or, in places, overlain with erosional unconformity by dark mid-cretaceous mudstones. distribution. central jameson land and south-eastern traill ø (figs 1, 2a, 2c). chronostratigraphy. upper callovian – lower middle oxfordian, including the p. athleta, q. mariae and c. den­ siplicatum ammonite zones, based on ammonites and dinoflagellate cysts. subdivision. the formation is subdivided (from below) into the athene, hades and zeus members. several jurassic lithostratigraphic units included here were named after geographical features in the area whose nomenclature followed a greek mythological theme, such as the mountains of olympen, pelion and parnas. no further place names were available in the region so that additional names from greek mythology are used to name the three constituent members of the olympen formation. key references. surlyk et al. (1973), birkelund et al. (1971), surlyk (1991), callomon (1993), larsen & surlyk (2003), bruhn & surlyk (2004), vosgerau et al. (2004b), bjerager et al. (2018a). athene member new member history. corresponds to the lower sandy unit of the olympen formation of surlyk et al. (1973). name. from athene (danish spelling of athena) who was the daughter of zeus, eponymous goddess of the city of athens, goddess of wisdom. type section. olympelven, central jameson land, 71°24.4′n, 23°34.8′w (larsen & surlyk 2003, fig. 2; figs 2a, 48). reference sections. mikael bjerg, parnas, pelion (fig. 2a). thickness. 57 m in the type section. lithology. massive, micaceous fineto medium-grained sandstones, forming sheet-like beds in the lower part and thick lenticular beds in the upper part, alternating with thin mudstones and heteroliths. fig 49 hades mb athene mb fig. 49 sandstones of the athene member (57 m thick) overlain by dark mudstones of the hades member, olympen formation, olympen, jameson land (figs 1, 2a). the section corresponds to the log shown in fig. 48. photograph by rikke bruhn. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 60 of 116 www.geusbul let in.org fossils. plant fragments, trace fossils, dinoflagellate cysts and very rare ammonites. depositional environment. slope and base-of-slope. boundaries. the lower boundary is defined at the base of the lowest massive micaceous sandstone, overlying silty mudstone of the fossilbjerget formation with a sharp boundary. the upper boundary is placed where the uppermost massive sandstone bed of the athene member is abruptly overlain by a uniform succession of dark silty mudstones referred to the hades member (fig. 49). distribution. known from outcrops in central jameson land, and from slim-core boreholes in central western jameson land in the lollandselv and falsterelv river valleys (bjerager et al. 2018a). the distribution in central western jameson land cannot be mapped in the field due to a cover of superficial deposits. chronostratigraphy. upper callovian p. athleta ammonite zone, and possibly q. lamberti ammonite zone, based on ammonites and indirectly on dinoflagellate cysts. includes faunal horizon j-37 of callomon (1993). hades member new member history. described as the middle muddy unit of the formation by surlyk et al. (1973). name. from hades who was the ruler of the underworld according to greek mythology. type section. olympelven, central jameson land, 71°24.4′n, 23°34.8′w (larsen & surlyk 2003, fig. 2; figs 2a, 48). thickness. 39 m in the type section. lithology. dark laminated mudstones with heterolithic intercalations. fossils. ammonites, bivalves, belemnites, dinoflagellate cysts, plants and trace fossils. depositional environment. base-of-slope and basin plain. boundaries.the lower boundary is placed at the base of dark laminated mudstones, overlying bioturbated, micaceous silty mudstones of the fossilbjerget formation in southern jameson land (see below under the zeus member), and massive sandstones of the athene member in central jameson land (fig. 49). the upper boundary is placed at the base of the lowest coarsegrained, cross-bedded or massive sandstone of the zeus member in central jameson land. several hundreds of metres north of the type section, slope gullies filled with massive sandstones are incised into the upper hades member (larsen & surlyk 2003, fig. 2). these sand bodies are assigned to the overlying zeus member and the upper boundary of the hades member is thus defined locally by the erosional base of these lenticular bodies. the member wedges out towards the south, and in southern jameson land it is represented by a dark grey mudstone succession, up to a few metres thick, overlain by black laminated mudstones with thick sandstone injectite bodies of the hareelv formation. distribution. central and southern jameson land. chronostratigraphy. lower middle oxfordian, including the q. mariae ammonite zone, based on ammonites and indirectly on dinoflagellate cysts. includes faunal horizon j-38 of callomon (1993). zeus member new member history. described as the upper sandy unit of the formation by surlyk et al. (1973). name. from zeus who was the ruler of the ancient greek olympic gods, the highest god in the greek pantheon. type section. olympelven, central jameson land, 71°24.4′n, 23°34.8′w (larsen & surlyk 2003, fig. 2; figs 2a, 48). thickness. at least 50 m, probably as much as 250 m. lithology. coarse-grained, cross-bedded sandstones, locally with wave ripples, and massive sandstones to the south. fossils. rare ammonites, plants and trace fossils. depositional environment. shallow marine shelf on traill ø, shelf-margin wedge and slope and base-of-slope in central jameson land. boundaries. the lower boundary is placed at the base of the lowest sandstone, abruptly overlying dark mudstones of the hades member in central jameson land, and overlying silty mudstones of the fossilbjerg formation with a gradational or sharp boundary on traill ø. the upper boundary is not exposed at the type section. elsewhere, it is placed at the base of black mudstones http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 61 of 116 www.geusbul let in.org with sandstone injectites of the hareelv formation in jameson land, and at the base of black mudstones of the bernbjerg formation on traill ø. distribution. central jameson land and south-eastern traill ø. in southern jameson land, thin distal hades member mudstones, overlying the fossilbjerget formation, coarsen upward into facies transitional to the zeus member. chronostratigraphy. lower–middle oxfordian, q. mariae – c. densiplicatum zones, based on ammonites and indirectly on dinoflagellate cysts. includes faunal horizons j-38 to j-40 of callomon (1993). 430 420 410 400 390 380 370 360 350 340 top pelion fm fo ss ilb je rg et f m h ar ee lv fm (p ar s) g on io m ya kl øf t m b cl si scl si s 530 520 510 500 490 480 470 460 450 440 560 570 m 550 540 450 m 440 h ar ee lv fm (p ar s) fig 50 fig. 50 type section of the hareelv formation, katedralen, ugleelv, jameson land (figs 1, 2a). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 62 of 116 www.geusbul let in.org hareelv formation history. erected by surlyk et al. (1973). in a provisional presentation of this revised stratigraphy, surlyk (2003, fig. 5) transferred the sjællandselv and salix dal members from the overlying raukelv formation to the hareelv formation and an additional member (katedralen member) was provisionally introduced. these proposals were followed in bjerager et al. (2018a, b) but are not formalised here. the hareelv formation is not subdivided, and the sjællandselv and salix dal members 560 550 540 530 520 510 500 490 480 470 460 450 440 h ar ee lv fm (p ar s) 430 420 410 400 390 380 370 fig 51 cl si scl si s m 600 590 580 570 fig. 51 reference section of the hareelv formation, gåseelv, jameson land (figs 1, 2a). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 63 of 116 www.geusbul let in.org are retained in the raukelv formation as originally defined by surlyk et al. (1973). type section. the original, rather poorly exposed type section of surlyk et al. (1973) is here substituted by a new, well-exposed type section at katedralen, southern jameson land, 70°53.3′n, 22°55.2′w (figs 1, 2a, 50). reference sections. katedralen, gåseelv (fig. 51), hareelv, goniomyakløft (surlyk et al. 1973, fig. 14; fig 2a); bloklev-1 slim-core, available for study at geus, copenhagen (see bjerager et al. 2018 a, b). thickness. probably up to about 400 m. lithology. dark laminated organic-rich mudstones interbedded with massive lenticular sandstone injectites. pervasive sand injectites, including dykes and sills on all scales (figs 50–52). mudstone clasts and thin finegrained turbidite sandstones are common. fossils. ammonites, bivalves, dinoflagellate cysts, logs, minute trace fossils at some levels. depositional environment. slope, base-of-slope and basin plain. boundaries. in south-west jameson land, the lower boundary is placed at the base of black, laminated mudstones, sharply overlying a few metres thick wedge of dark hades member mudstones of the olympen formation. elsewhere, it is placed where black laminated mudstones overlie sandstones of the zeus member (olympen formation) or grey micaceous silty mudstones of the fossilbjerget formation. the hareelv formation is typically succeeded conformably by the raukelv formation, the boundary being placed at the base of marine, massive sheet-like sandstones of the sjællandselv member. it is possible that the formation is overlain by the clinoform-bedded langelandselv member in the subsurface of south-west jameson land. in a localised area of southernmost jameson land, the upper boundary is placed at a sharp, erosional surface overlain by bioturbated mudstones referred to the lower cretaceous (ryazanian) hesteelv formation. further details on this boundary are given under the hesteelv formation. distribution. southern jameson land (figs 1, 2a). chronostratigraphy. upper oxfordian – lower volgian, a. glosense (faunal horizon j-41 of callomon, 1993) – p. ele­ gans zones, based on ammonites and dinoflagellate cysts. in the subsurface (south-west jameson land), the formation extends down into the middle oxfordian c. den­ siplicatum zone (bjerager et al. 2018a). upper boundary diachronous, younging towards the south. key references. surlyk et al. (1973), surlyk (1987), surlyk & noe-nygaard (2001b, 2003), surlyk et al. (2007), alsen & piasecki (2018), bjerager et al. (2018b). fig. 52 hareelv formation, showing black mudstones and complex sandstone injectites. section about 50 m high. katedralen, ugleelv, southern jameson land (figs 1, 2a). http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 64 of 116 www.geusbul let in.org bernbjerg formation history. erected by surlyk (1977b). type section. bernbjerg, kuhn ø 74°46.9′n, 20°19.7′w (surlyk 1977a, fig. 24, section 8; figs 1, 2e, 53). reference sections. the reference sections follow the numbering of surlyk (1977a) and are illustrated therein: bernbjerg (section 9), perisphinctes ravine (section 10), ‘nord profil’ kuhn ø (section 12), cardiocerasdal (sections 38, 40), stratumbjerg (section 41; fig. 54; figs 1, 2e). thickness. up to about 600 m. lithology. dark grey to black mudstones, which may be silty or sandy. basal 20–75 m dominated by mudstone 10 36 m poorly exposed 8 m poorly exposed 41 m poorly exposed 43 m poorly exposed 50 60 70 80 130 140 150 270 260 250 240 230 180 170 160 270 280 fig 53 290 300 320 380 m 370 360 350 330 340 m 310 m m 0 cl si s cl si s cl si s cl si s be rn bj er g fm (p ar s) pe lio n fm u gp ik r av in e m b fig. 53 type section of the bernbjerg formation, bernbjerg, kuhn ø (figs 1, 2e). from surlyk (1977a, fig. 24). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 65 of 116 www.geusbul let in.org or sandy heteroliths. thin graded sandstone turbidites are common in the mudstones in some areas. fossils. ammonites, belemnites, the bivalve buchia, wood fragments and logs, rare vertebrates, dinoflagellate cysts and trace fossils. depositional environment. offshore marine. boundaries. the lower boundary is placed where dark grey to black mudstones and sandy heteroliths abruptly overlie sandstones of the olympen formation (traill ø), sandstones of the payer dal formation (hold with hope, kuhn ø, hochstetter forland and store koldewey), or sandstones or heteroliths of the jakobsstigen formation (wollaston forland). the lower boundary is interpreted as a regional, northwards younging, diachronous drowning surface. the upper boundary is an erosional unconformity at the base of the conglomerate-dominated rigi and laugeites ravine members of the lindemans bugt formation and, more distally (eastwards) in wollaston forland by dark sandy mudstones of the niesen member (wollaston forland − kuhn ø). on traill ø and store koldewey, it is overlain with erosional unconformity by aptian mudstones of the stratumbjerg formation. distribution. traill ø, hold with hope, east coast of clavering ø, wollaston forland, kuhn ø, th. thomsen land, hochstetter forland, store koldewey (figs 1, 2c, 2d, 2e). chronostratigraphy. upper oxfordian – lower volgian, a. glosense – probably p. wheatleyensis ammonite zones, based on ammonites. subdivision. a new member, the ugpik ravine member, is defined here. key references. maync (1947), sykes & surlyk (1976), surlyk (1977a, 1978c, 2003), alsgaard et al. (2003), vosgerau et al. (2004a, b), bjerager et al. (2020). fig 54 0 10 20 30 40 50 60 70 80 90 m m s f m c cl/si scl/si f m c be rn bj er g fm pa ln at ok es b je rg f m u gp ik r av in e m b (p ar s) rø dr yg ge n m b (p ar s) slight angular and erosional unconformity fig. 54 reference section of the bernbjerg formation, stratumbjerg, wollaston forland (figs 1, 2e). from surlyk (1977a, fig. 32). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 66 of 116 www.geusbul let in.org ugpik ravine member new member history. the member includes some of the deposits described under the term ‘grey series’ by maync (1947). the ‘grey series’, however, referred to a heterolithic facies type that occurs at several levels in the middle– upper jurassic succession of wollaston forland and kuhn ø (maync 1947; surlyk 1977a). the ‘grey series’ is thus a facies term and not a lithostratigraphic unit and did not form the basis for a formal lithostratigraphic unit in the scheme of surlyk (1977a). name. after a creek in eastern payer dal, kuhn ø, termed ugpik ravine by maync (1947, p. 18). ugpik is greenlandic for snowy owl. type section. ugpik ravine, southern kuhn ø, 74°45.3′n, 20°21.3′w (surlyk 1977a, fig. 4, section 1; figs 1, 2e, 55). reference sections. the reference sections follow the numbering of surlyk (1977a) and are illustrated therein: wollaston forland (section 38; fig. 46) and kuhn ø (sections 1, 8, 9; figs 1, 2e). thickness. 85–100 m in cardiocerasdal, wollaston forland, and up to about 75 m on kuhn ø (figs 1, 2e). lithology. bioturbated, mainly wavy laminated sandy heteroliths, but flaser and lenticular laminated heteroliths also occur. fossils. ammonites, belemnites, bivalves, plant fragments, dinoflagellate cysts and trace fossils. 60 50 40 30 20 10 0 130 140 150 160 170 120 110 100 90 80 70 m m m s f m c cl/si scl/si scl/si f m c f m c u gp ik r av in e m b be rn bj er g fm u gp ik r av in e m bbe rn bj er g fm (p ar s)pa ye r d al fm (p ar s) fig 55 fig. 55 type section of the ugpik ravine member, ugpik ravine, kuhn ø (figs 1, 2e). from surlyk (1977a, fig. 4). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 67 of 116 www.geusbul let in.org depositional environment. offshore transition zone. boundaries. the lower boundary is the same as for the formation, and is placed at a marked drowning surface, separating the basal heteroliths of the member from the underlying sandstones of the payer dal formation on hold with hope, kuhn ø, hochstetter forland and store koldewey, and from the underlying sandstones or heteroliths of the jakobsstigen formation on wollaston forland. the upper boundary is sharp and placed at the base of mud-dominated heteroliths or laminated mudstones of the upper bernbjerg formation, overlying sandy heteroliths. this boundary represents a drowning surface (figs 45, 46, 55). distribution. hold with hope, wollaston forland, th. thomsen land, kuhn ø and hochstetter forland (figs 1, 2e). chronostratigraphy. upper oxfordian (a. glosense, a. ser­ ratum, a. regulare and a. rosenkrantzi ammonite zones) – lower kimmeridgian (p. baylei and r. cymodoce zones), based on ammonites. raukelv formation revised formation introductory remarks. the raukelv formation was erected by surlyk et al. (1973) for a succession of coarse-grained, marine high-angle clinoform-bedded, large-scale cross-bedded or massive sandstones of volgian age exposed in southernmost jameson land. it represents the final infilling of the basin, forming the regressive upper part of the middle jurassic – lowermost cretaceous transgressive–regressive megacycle. the massive sandstones of the sjællandselv member and the mudstones of the salix dal member represent slope–basinal redeposited sands and basinal muds, respectively. the sjællandselv member represents basinal sedimentation during each progradational clinoform pulse formed by collapse of the clinoform fronts, and the salix dal member represents basinal mud accumulation during each intervening shelf-edge aggradational phase. the raukelv formation is very well exposed but due to the combination of low altitude and subdued topography, and the great thickness of the clinoform beds, many sections only show a single bed and long, continuous vertical sections through a number of stacked beds can only rarely be measured. some of the figured type or reference sections therefore only show part of the unit in question, and the uppermost member – the rauk plateau member – is only illustrated with field photographs. history. as originally defined by surlyk et al. (1973), the raukelv formation comprised three members, in stratigraphic order, the sjællandselv, salix dal and fynselv members. the fynselv member is revised here, being restricted to include only the major, landscape-forming high-angle clinoform bed, which formed the main element (bed d) of the original definition of the member by surlyk et al. (1973). name. after the raukelv river in southernmost jameson land (figs 1, 2a). type area. plateau around the upper reaches of the raukelv river (fig. 2a). as noted above, the formation is defined from a number of short sections that typically exhibit only a single member or part thereof. the formation is thus defined by the type sections of its component members. reference sections. good sections are found in canyons of the eastern and western branches of the fynselv river (fig. 2a), and in the northern reaches of the langelandselv river (fig. 2a). thickness. probably around 400 m, but difficult to estimate as only a few members are exposed in even the largest vertical outcrops due to the low altitude and hence subdued relief of southernmost jameson land. lithology. high-angle clinoform-bedded, coarse-grained, commonly pebbly sandstone beds, up to 50 m thick, alternate with intervals of large-scale planar or medium scale trough cross-bedded pebbly sandstones and bioturbated siltstones. in some cases, the upper, frontal part of thick clinoform beds display curved slump scars draped by massive sandstones. similar massive resedimented sandstones forming parallel-bedded successions are referred to the sjællandselv member. such massive sandstone beds are very uniform, mainly unfossiliferous, and are essentially devoid of any diagnostic features. fossils. ammonites, bivalves, belemnites, crinoids, trace fossils, wood and plant fragments. depositional environment. extensive shelfal sandwave fields, shelf-margin wedges, shelf-slope break, slope, base-of-slope and basin. boundaries. the lower boundary is placed at the base of the lowest thick, massive sandstone bed of the sjællandselv member, overlying black mudstones with sandstone injectites of the hareelv formation. the upper boundary is placed at a major canyon-shaped erosional http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 68 of 116 www.geusbul let in.org sb4 150 m a facies members key surfaces o-sb4 o-sb3 o-sb2 o-ss3 o-ss1 langelandselv mb prasino mb fynselv mb kokkino mb straight river mb rauk plateau mb o-ss2 o-sd3 o-sd2 o-sd1 sb3 sb2 ss2 ss1 sb1 fig 56a_version 2 ss3 si si sd3 sd1 sd2100 50 0 raukelv fm (lower – upper volgian) hareelv fm (upper oxfordian – upper volgian) hesteelv fm (lower ryazanian – lower valanginian) sjællandselv mb (lower–upper volgian) fynselv mb (middle volgian) straight river mb black mudstoneshelf-margin wedge with high-angle clinoform bedding shelf-margin wedge with low-angle clinoform bedding massive sandstone bodies formation boundary member boundary fig 56b_version 3 crinoid bjerg mb (lower ryazanian) base ryazanian unconformity rauk plateau mb (upper volgian) muslingeelv mb (lower ryazanian – lower valanginian) n s b fig. 56 raukelv formation stratigraphic relationships. (a) schematic section through the upper part of the raukelv formation, above the sjællandselv and salix dal members, showing symbols used for beds and surfaces. sb: sand bank. sd: sand dune. ss: sand sheet. si: siltstone. o-sb, o-sd and o-ss indicate omission surfaces bounding the individual beds. scale is approximate. this figure is modified from surlyk & noe-nygaard (1991, fig. 3); subsequent work has shown that the lower and upper si units pass laterally into clinoform-bedded sandstones of the fynselv and straight river members, respectively (see text for further discussion). (b) schematic n–s stratigraphic section showing the inferred stratigraphic relationships between the shelf and slope/basin components of the raukelv formation and associated formations. note the distal, shelf-edge position of the rauk plateau member. modified after surlyk & noe-nygaard (2005, fig. 3). http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 69 of 116 www.geusbul let in.org unconformity, where high-angle clinoform-bedded, cross-bedded pebbly and massive sandstones of the rauk plateau member are overlain by mudstones and siltstones of the crinoid bjerg member (hesteelv formation), and higher up the erosional relief by richly fossiliferous pebbly sandstones and massive sandstones of the muslingeelv member at the top of the hesteelv formation. distribution. the raukelv formation forms the top plateau in southernmost jameson land (figs 1, 2a). chronostratigraphy. volgian. the lower part of the formation has yielded lower volgian ammonites and the bivalve buchia mosquensis. the bulk of the formation is of middle volgian age and contains abundant ammonites. erosional remnants below the top unconformity contain rare upper volgian ammonites. subdivision. the formation is subdivided (from below) into the sjællandselv, salix dal, langelandselv, prasino, fynselv, kokkino, straight river and rauk plateau members. as noted under the hareelv formation, the provisional transfer of the sjællandselv and salix dal members by surlyk (2003) is not formalised here. these members are retained in the raukelv formation as originally defined by surlyk et al. (1973). with the exception of the sjællandselv and salix dal members at the base of the formation, the remaining members relate to the subdivision of surlyk & noenygaard (1991) based on a facies classification and sequential key surfaces. the main facies types – sb sand bank, sd sand dunes, ss sand sheet and si siltstone – were numbered in ascending order as e.g. sb1, sb2…, sd1, sd2…. etc., and their tops, which are omission surfaces, as o-sb1, o-sb2…. etc. (fig. 56a). the latter figure (surlyk & noe-nygaard 1991, fig. 3) is reproduced here since this depicted the stratigraphic relationships that formed the basis for the original definitions. subsequent work has shown, however, that the silty mudstone (si) units are less persistent than depicted and are only locally developed, representing bottomsets and basinal facies deposited in front of prograding clinoforms (fig. 56b). key references. aldinger (1935), surlyk et al. (1973, 1975, 1993, 2003), surlyk & noe-nygaard (1991, 1995, 2005). sjællandselv member revised member history. this unit was erected by surlyk et al. (1973) as the lower member of the raukelv formation. type section. sjællandselv, 70°42.8′n, 23°23.3′w (surlyk et al. 1973, fig. 31b, section 14; figs 2a, 57). thickness. probably up to 100 m thick. lithology. massive sandstones, typically sheet-like, with scattered mudstone clasts up to about 0.5 m in length. thin mudstone interbeds occur in the lower part. fossils. rare ammonites and bivalves. depositional environment. slope, base-of-slope and basin in front of clinoform-bedded shelf-margin wedges. the slope sandstones interdigitate with the shelf-margin wedges (figs 56b, 58). boundaries. the boundaries of the sjællandselv member are highly complex as the member represents a fringe of structureless gravity-flow sands shed from the collapsed front of successive shelf-margin clinoform beds si s ra uk el v fm hareelv fm langelandselv mb 0 20 40 60 80 100 120 140 160 m sa lix d al m b sjæ lla nd se lv m b fig 57 fig. 57 type sections of the sjællandselv and salix dal members, raukelv formation, salix dal, jameson land (figs 1, 2a). modified from surlyk et al. (1973, fig. 31b). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 70 of 116 www.geusbul let in.org of the raukelv formation (fig. 56b). the lower boundary is placed where the sandstone/mudstone ratio changes from about 50% in the hareelv formation to about 90% in the sjællandselv member, where the sand injectites characterising the hareelv formation disappear, and where the extremely irregular outcrop topography and lack of bedded sandstones in the hareelv formation are replaced by sheet-like sandstones of the sjællandselv member. upslope, the sjællandselv member interfingers laterally with shelf-margin clinoform beds (figs 56b, 58). the upper boundary in the type section is placed where dark mudstones referred to the salix dal member abruptly overlie parallel-bedded massive sandstones. in southern jameson land, this boundary is strongly diachronous and is placed where the member underlies or onlaps the front of the clinoform-bedded members of the raukelv formation. in the extreme south-eastern outcrop area, the sjællandselv member is locally overlain, at an erosional unconformity, by the hesteelv formation. distribution. southern jameson land. chronostratigraphy. volgian, based on ammonites and dinoflagellate cysts. salix dal member revised member history. erected by surlyk et al. (1973) as the second member from below of the raukelv formation. type section. sjællandselv, southern jameson land, 70°40.8′n, 23°23.4′w (surlyk et al. 1973, fig. 31b, section 14; figs 1, 2a, 57). thickness. up to about 80 m, wedges out towards the north and north-east. lithology. dark micaceous silty and sandy mudstones with mainly thin sandy turbidites. slump folds ubiquitous. fossils. rare ammonites, dinoflagellate cysts. depositional environment. slope, base-of-slope and basin. boundaries. the lower boundary is placed at the base of dark mudstones, overlying massive sandstones of the sjællandselv member. in the type section, the upper boundary is placed at the base of the lowermost coarsegrained clinoform or cross-bedded sandstone bed of the langelandselv member. the salix dal member is interpreted to be succeeded by younger sandstone members of the raukelv formation farther basinward, but such relationships are not observed at outcrop. distribution. south-western jameson land (fig. 1). chronostratigraphy. volgian, based on ammonites and dinoflagellate cysts. fig. 58 person indicating the boundary between well-structured sandstones of the fynselv member to the left and structureless, slumped sandstones to the right. the structureless sandstones represent the up-slope feather-edge of the basinal sjællandselv member sandstones, formed by collapse of the clinoform front at the shelf edge. nw-oriented branch of the muslingeelv river, southern jameson land (figs 1, 2a). http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 71 of 116 www.geusbul let in.org sb2 0 5 10 15 20 sb1 0-sb1 ss1 fig 59 ra uk el v fm la ng el an ds el v m b (p ar s) pr as in o m b ra uk el v fm fy ns el v m b pr as in o m b 45 50 55 60 65 70 73 m 0-ss2 ss2 0-sd1 0-ss1 sd1 25 30 35 m 37 40 poorly exposed ssicl f m c ssicl f m c fig. 59 type sections of the langelandselv and prasino members and reference section of the fynselv member, langelandselv, sw jameson land (figs 1, 2a). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 72 of 116 www.geusbul let in.org langelandselv member new member name. after the langelandselv river in southernmost jameson land, where the member is well exposed (fig. 2a). type section. west side of the upper reaches of the langelandselv river, south-west jameson land, 70°35.2′n, 23°23.0′w (figs 1, 2a, 59). reference section. west side of the upper reaches of the langelandselv river, a few hundred metres north of the type section (fig. 60). m fig 60 11 12 13 14 15 16 17 18 19 20 21 22 1 0 2 3 4 5 6 7 8 9 10 ss1 11 m s pb vf sicl f m c vc vf f m c vc s pb vf sicl f m c vc vf f m c vc o-ss1 sd1 o-sd1 ss2 pr as in o m b (p ar s) la ng el an ds el v m b ra uk el v fm la ng el an ds el v m b (p ar s) ra uk el v fm fig. 60 reference sections of the langelandselv and prasino members. west side of the upper reaches of the langelandselv river, southern jameson land (figs 1, 2a). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 73 of 116 www.geusbul let in.org thickness. maximum thickness is about 20 m. lithology. coarse-grained, commonly pebbly sandstone. the basal unit, a high-angle clinoform bed (sb1) with cross-bedded intrasets and abundant internal truncation surfaces, is overlain by a trough cross-bedded, strongly burrowed sandstone unit (ss1), capped by a regional transgressive surface of erosion (o-ss1; figs 56a, 59, 60). fossils. ammonites, bivalves, crinoids, trace fossils, plant and wood fragments. depositional environment. south-eastward prograding shelf-margin wedge followed by shallow marine shoreface. boundaries. the lower boundary is not exposed at outcrop but the member is inferred to overlie massive sandstones with minor thin, black mudstones of the sjællandselv member or black mudstones with thin sandstones of the salix dal member. the upper boundary is placed at a marked omission surface (o-ss1) at the base of the lowermost large-scale cross-bedded sandstone bed (sd1) of the prasino member. distribution. probably the same as the formation but the member is only exposed in the western part of the outcrop area of the formation. chronostratigraphy. lower – lower middle volgian, based on ammonites and bivalves. prasino member new member name. after the greek word for ‘green’, reflecting the colour, especially of the upper part of the member in many outcrops. type section. west side of the upper reaches of the langelandselv river, same as type section of the langelandselv member, 70°35.2′n, 23°23.0′w (figs 2a, 59). reference sections. western branch of the fynselv river (fig. 61) and west flank of the langelandselv river (figs 2a, 60). thickness. maximum thickness estimated at about 25 m. lithology. the member comprises a lower unit of coarsegrained, large-scale cross-bedded sandstone (sd1), capped by an omission surface (o-sd1; figs 62–64), o-sb2 sb2 ss2 o-ss2 fig 61 s pb vf sicl f m c vc vf f m c vc0 5 10 15 20 25 25 m pr as in o m b (p ar s) fy ns el v m b (p ar s) ra uk el v fm fig. 61 reference sections of the prasino and fynselv members. west branch of the fynselv river, southern jameson land (figs 1, 2a). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 74 of 116 www.geusbul let in.org fig 62 n o-ss2 o-ss2 o-sd1 sd1 o-sd1 o-sb1 sb1 sd1 ss2 o-sd1 o-sd1 sd1 ss2 104114 96 104 110 96 s ns fig. 62 line drawings from field photographs of the prasino member, showing the large-scale cross-bedded lower part of the prasino member (sd1) overlain by low-angle trough cross-bedded and strongly burrowed bed (ss2), raukelv formation (see fig. 56). plateau east of the western branch of the fynselv river, southern jameson land (figs 1, 2a). the upper transect is modified after surlyk & noe-nygaard (1991, fig 13). fig. 63 large-scale cross-bedded lower part of the prasino member (sd1) overlain by low-angle trough cross-bedded and strongly burrowed bed (ss2), raukelv formation (see fig. 56). plateau east of the western branch of the fynselv river, southern jameson land (figs 1, 2a). fig 63 ss2 sd1 o-sd1 fig 64 ss2 o-sd1 sd1 fig. 64 close-up of section shown in fig. 63 showing large-scale cross-bedded sandstones (sd1) with tidal bundles, overlain by trough cross-bedded and burrowed sandstones (ss2) of the prasino member, raukelv formation (see fig. 56). plateau east of the western branch of the fynselv river, southern jameson land land (figs 1, 2a). http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 75 of 116 www.geusbul let in.org overlain by greenish, coarse-grained, commonly pebbly sandstone showing poorly preserved trough cross-lamination (ss2), capped by an omission surface (o-ss2). this part of the member (ss2) is a highly variable, mainly loose and poorly exposed unit composed of coarsegrained, trough cross-bedded sandstone, commonly green in colour (this facies is not shown in surlyk & noenygaard 1991, fig. 3; fig. 56a). fossils. ammonites, bivalves, logs and trace fossils, including the metre-long u-burrows of tisoa habichi, mainly in the upper part. depositional environment. shallow marine, tidal sandwave fields and offshore marine. boundaries. the lower boundary (o-ss1) is placed at the base of large-scale trough cross-bedded sandstone of sd1, overlying mainly cross-laminated, strongly bioturbated and diffusely bedded sandstone of ss1. the upper boundary is an important omission surface (o-ss2), placed at the base of a thick high-angle clinoform-bedded sandstone (sb2) or at the base of a siltstone unit (si) that in a few areas forms the lower part of a coarsening-upwards succession passing upwards into the clinoform bed of sb2 (fig. 59). distribution. exposed in the northern outcrop areas of the formation but appears to have wedged out towards the east. chronostratigraphy. middle volgian. fynselv member revised member history. the member is revised here from that defined by surlyk et al. (1973) to include only the main plateau-forming lower clinoform bed of the original fynselv member (bed d of surlyk et al. 1973, denoted sb2 in surlyk & noe-nygaard 1991; fig 56a). the lower part of the original fynselv member is here referred to the new langelandselv and prasino members. the upper part of the original member is here referred to the new kokkino, straight river and rauk plateau members. type area. eastern branch of the fynselv river and the adjacent western plateau, southern jameson land (figs 1, 2a). reference sections. western branch of the upper reach of langelandselv (fig. 2a); same as type sections for the 0 5 10 15 20 25 m ssicl fvf m c vc pb sb2 o-ss2 ss2 fig 65 pr as in o m b (p ar s) fy ns el v m b (p ar s) ra uk el v fm fig. 65 reference section of the fynselv member (sb2 in surlyk & noe-nygaard 1991). eastern branch of the fynselv river, jameson land (figs 1, 2a). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 76 of 116 www.geusbul let in.org langelandselv and prasino members (fig. 59), west bank of fynselv (figs 2a, 61, 65). thickness. maximum thickness about 50 m, but difficult to estimate as the member forms the top plateau in the main outcrop area (fig. 66). lithology. very coarse-grained, pebbly, plateau-forming sandstone showing high-angle (~25°) mainly sigmoidal clinoform bedding (sb2, in surlyk & noe-nygaard 1991; figs 56a, 67). it differs from the other high-angle clinoform-bedded members in its greater thickness, uniformity, great lateral extent (mapped over 900 km²), abundance of bivalves, and especially in the regular, mainly sigmoidal nature of the clinoforms, in some cases with a relatively thick topset. although the clinoforms of the fynselv member may locally be truncated and tangential, the other clinoform-bedded members characteristically show tangential clinoforms throughout, commonly with internal erosional reactivation surfaces. slump scars, draped by massive sandstones, are present locally in the fynselv member (fig. 58). fossils. ammonites (pavlovia, dorsoplanites), bivalves (camptonectes, entolium, astarte, isocyprina), crinoids, wood, large plant fragments and trace fossils, including curvolithos and tisoa habichi. depositional environment. eastward prograding shelf-margin wedge. fig. 66 the plateau-forming high-angle clinoform bed of the fynselv member (c. 30 m thick), showing a typical example of the outcrop of this landscape-forming unit. this bed can be traced over 900 km2 and reaches thicknesses of up to 50 m. western branch of the fynselv river, southern jameson land (figs 1, 2a), viewed looking west. fig 67 o-sb2 fig. 67 the fynselv member showing large-scale clinoform bedding with asymptotic topsets creating sigmoidal clinothems, capped by an omission surface (o-sb2, see fig. 56), raukelv formation, southern jameson land (figs 1, 2a). person (encircled) for scale. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 77 of 116 www.geusbul let in.org boundaries. the lower boundary is placed at the base of a thick, laterally extensive high-angle clinoform-bedded sandstone (sb2), overlying greenish, very coarse-grained pebbly sandstones of the upper prasino member (ss2) at a sharp surface (o-ss2). in western outcrops, in the langelandselv area, the boundary is placed at the sharp base (o-ss2) of an up to 14 m thick strongly burrowed siltstone which coarsens gradationally upwards into the clinoform bed of sb2 (fig. 59; lower ‘si’ unit in fig. 56a). the upper boundary is placed at the base of the lowest strongly bioturbated and cross-bedded sandstone (ss3) at a marked omission surface (o-sb2), forming the topographic top surface of the main central plateau in the outcrop area (figs 56a, 67). distribution. same as the formation except for the eastern areas where it has been removed by modern erosion. chronostratigraphy. middle volgian, based on ammonites and bivalves. kokkino member new member history. the member was included in the upper part of the fynselv member of surlyk et al. (1973) and corresponds to the ss3, sd2 and sb3 units of surlyk & noenygaard (1991) (fig. 56a). name. after the greek word for ‘red’, reflecting the commonly red colour of the top part of the member in outcrop. type section. ridge on the east side of the eastern branch of the fynselv river (figs 2a, 68), position somewhat uncertain. thickness. maximum thickness estimated at about 40 m. lithology. medium and coarse-grained, trough cross-bedded burrowed, commonly pebbly, sandstones (ss3), overlain by cross-bedded tidally influenced sandstones (sd2), and high-angle clinoform-bedded sandstone (sb3). the clinoforms are tangential or rarely sigmoidal and the clinothems locally show large-scale trough cross-bedded intrasets (fig. 56a). fossils. trace fossils. depositional environment. shelfal sandwave fields followed by an east to south-eastward prograding clinoform-bedded shelf-margin wedge. boundaries. the lower boundary is placed at the base of the lowest cross-bedded and strongly bioturbated sandstones (ss3), overlying the omission surface (o-sb2) capping the clinoform-bedded fynselv member (sb2). the upper boundary is placed at the base of low-angle sb3 o-sd2 o-sb2 o-ss3 sd2 sb2 ss3 s vf f m c vc pb fig 68 0 5 10 15 20 m ko kk in o m b (p ar s) fy ns el v m b (p ar s) ra uk el v fm fig. 68 type section of the kokkino member (ss3, sd2, sb3 in surlyk & noe-nygaard 1991, see fig. 56). the log shows only the gently inclined bottomset of the steep clinoform bed of sb3. eastern branch of the fynselv river, jameson land (figs 1, 2a). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 78 of 116 www.geusbul let in.org clinoform-bedded sandstones referred to the straight river member that abruptly overlie the high-angle clinoform-bedded sandstones of the kokkino member at the o-sb3 omission surface (fig. 56a). in a few areas, the clinoform-bedded sandstones of the straight river member are underlain by a siltstone unit (upper si unit in fig. 56a) in such cases, the upper boundary is placed at the base of the siltstone, overlying the clinoform bed of the kokkino member at the omission surface o-sb3 (fig. 56a). distribution. probably the same as the formation, but the member has been removed by modern erosion at the western outcrop margins. chronostratigraphy. probably middle volgian. straight river member new member history. the member was not known at the time when the raukelv formation was erected, and only in part (sd3) at the time when surlyk & noe-nygaard (1991) was published. it was thus only implicitly included in the upper part of the fynselv member of surlyk et al. (1973). the part labelled sd3 in fig. 56a is thus much thicker and replaces the underlying si unit in most areas of exposure of the member. name. from straight river (fig. 2a), named by aldinger (1935). type section. eastern side of straight river, southern jameson land, 70°30.0′n, 23°07.8′w (figs 1, 2a, 69). reference section. southern part of straight river (figs 2a, 70). thickness. maximum exposed thickness estimated at about 50 m. the base and top of the member are nowhere exposed in the same sections. the clinoform set may thus attain much greater thicknesses than indicated on fig. 56a. lithology. coarse-grained sandstones, showing low angle (up to 7°) clinoform bedding showing large-scale cross-bedded, tidally influenced intrasets (sd3). in a few areas, the lower part of the member is composed of siltstones that pass upwards into the clinoform bed (upper si in fig. 56a). fossils. belemnites and trace fossils. 0 5 10 15 sd3 20 25 30 m s f m c sicl pb st ra igh t r ive r m b (p ar s) ra uk el v fm fig 69 fig. 69 type section of the straight river member, straight river, jameson land (figs 1, 2a). the large-scale cross-bedded sets are intrasets in low-angle clinothems. for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 79 of 116 www.geusbul let in.org depositional environment. south-eastward prograding shelf-margin wedge. boundaries. the lower boundary is placed at the base of the low-angle clinoform-bedded, sandstone unit with large-scale cross-bedded intrasets, or locally at the base of a siltstone unit, abruptly overlying the high-angle clinoform bed of the kokkino member at a marked omission surface (o-sb3). the upper boundary (o-sd3) is placed at the base of the high-angle clinoform bed of the rauk plateau member (sb4; fig. 56a), or at the angular and erosional unconformity at the base of the ryazanian hesteelv formation. distribution. the member is restricted to the eastern part of the outcrop area of the formation and has wedged out west of fynselv (fig. 2a). chronostratigraphy. probably lower upper volgian, as constrained by underlying and overlying members. rauk plateau member new member history. sandstones of this member were first described by aldinger (1935). the member formed the uppermost part of the fynselv member as defined by surlyk et al. (1973). described as sb4 by surlyk & noe-nygaard (1991). name. after the rauk plateau of aldinger (1935, plate 2). type section. between straight river and the raukelv river, 70°29.3′n, 23°09.2′w (figs 2a, 71, 72). no logs are presented as the member only comprises one high-angle clinoform bed exposed in vertical walls. thickness. maximum measured thickness is about 27 m (surlyk & noe-nygaard 2005). lithology. coarse-grained, pebbly high-angle (~25°) clinoform-bedded sandstones with large-scale cross-bedded intrasets and internal large-scale erosional reactivation surfaces (sb4; figs 56a, 72). the clinoforms are truncated-tangential or rarely sigmoidal in areas where the top is not truncated by the base-hesteelv formation unconformity. fossils. trace fossils, wood fragments. depositional environment. north-eastward prograding shelf-margin wedge. boundaries. the lower boundary is placed at an abrupt omission surface (o-sd3) that forms the base of the high-angle clinoform bed (sb4), overlying the low-angle, internally cross-bedded clinoform bed (sd3) of the straight river member. outwith the outcrop area of the hesteelv formation, the upper boundary of the rauk plateau member is typically the present-day top erosion surface. in a few places, however, the depositional upper boundary is preserved as an intensely bioturbated omission surface (o-sb4). in the outcrop area of the hesteelv formation, the upper boundary is an angular, erosional unconformity overlain by black to grey mudstones of the crinoid bjerg member or, to the north, by the sandy ra uk el v fm st ra igh t r ive r m b (p ar s) sd3 fig 70 0 5 10 15 20 m s f m c sicl pb fig. 70 reference section of the straight river member, straight river, jameson land (figs 1, 2a). the large-scale cross-bedded sets are intrasets in low-angle clinothems. for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 80 of 116 www.geusbul let in.org shell conglomerate of the lowermost muslingeelv member, both belonging to the hesteelv formation). distribution. the member is restricted to the southern and south-eastern outcrop areas of the raukelv formation, but its distribution area is probably greatly reduced due to modern erosion. chronostratigraphy. upper volgian, constrained by ammonites and bivalves in underlying and overlying units. fig. 71 photograph of the type section of the rauk plateau member, rauk plateau, jameson land, viewed towards the sse (figs 1, 2a). the underlying beds are rarely shown in the same section as the member, which forms the uppermost bed in most of the area. the type section is thus shown in a photograph. towards the east and south, the member is truncated by an erosional canyon filled in with the hesteelv formation. person (lower left) for scale. fig 72 fig. 72 photograph of the high-angle clinoform-bedded rauk plateau member. note the internal truncation surface downlapped by younger clinoforms (arrowed). immediately north of the type section, rauk plateau, southern jameson land (figs 1, 2a); person (lower right) for scale. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 81 of 116 www.geusbul let in.org hesteelv formation history. erected by surlyk et al. (1973) and described in more detail by surlyk (1973). type section. between the muslingeelv and hesteelv rivers, southern jameson land, 70°29.0′n, 23°02.3′w (surlyk et al. 1973, fig. 36a; surlyk 1973, fig. 7, section 6; figs 1, 2a, 73a). thickness. 10−120 m. lithology. the formation fills a c. 10 km wide u-shaped canyon with a southward plunging axis (fig. 74). black mudstones occur in the deepest part of the canyon and pass upwards into silty and sandy micaceous mudstones capped by a sandy shell conglomerate that, in the central part of the canyon, is overlain by a thin mudstone passing upwards into coarse-grained sandstone. fossils. ammonites, belemnites, bivalves, crinoids, wood and trace fossils. depositional environment. prograding shelf-margin canyon fill. boundaries. the lower boundary is placed at the base of dark mudstones of the crinoid bjerg member, overlying yellow sandstones of the rauk plateau or sjællandselv members (both of the raukelv formation), or black, well-laminated mudstones of the hareelv formation in a small area to the south-east (fig. 73a). the boundary is an angular, erosional unconformity, representing the floor of the u-shaped canyon. the top of the formation is the modern erosion surface. 0 cl si s 20 40 60 80 100 m us lin ge el v m b (p ar s) c rin oi d bj er g m b h es te el v fm h es te el v fm hareelv fm 120 0 20 40 60 80 100 a b m m m us lin ge el v m b (p ar s) c rin oi d bj er g m b (p ar s) fig 73 cl si s fig. 73 hesteelv formation, hesteelv, southernmost jameson land (figs 1, 2a). for legend, see fig. 7. (a) type sections of the hesteelv formation and its constituent crinoid bjerg and muslingeelv members. (surlyk et al. 1973, fig. 36a). (b) reference section of the muslingeelv member (surlyk et al. 1973, fig. 36b). fig. 74 contoured block diagram of a major erosional canyon incised in the se margin of the prograding rauk plateau member and subsequently filled with the upwards-coarsening hesteelv formation. several fluvial tributaries lead to the margin of the canyon. rauk plateau, southernmost jameson land (figs 1, 2a). http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 82 of 116 www.geusbul let in.org distribution. southernmost jameson land (figs 1, 2a). chronostratigraphy. ryazanian (lowermost cretaceous), praetollia maynci and hectoroceras kochi ammonite zones. subdivision. subdivided from below into the crinoid bjerg and muslingeelv members. key references. aldinger (1935), spath (1947), surlyk et al. (1973), surlyk (1973, 2003), surlyk & noe-nygaard (2005). crinoid bjerg member history. erected by surlyk et al. (1973) and described in more detail by surlyk (1973). type section. north-eastern slope of the crinoid bjerg mountain, 70°29.7′n, 23°04.0′w (figs 2a, 73a). thickness. maximum measured thickness of 85 m in the deepest part of the canyon, wedging out at the margins. lithology. black laminated mudstones in the deepest part of the canyon which cuts through the raukelv formation into the top of the hareelv formation (fig. 73a). this facies grades upwards and laterally into silty and sandy mudstones. fossils. ammonites, belemnites, crinoids, wood and plant debris, trace fossils. depositional environment. prograding shelf-margin canyon fill. boundaries. lower boundary same as the formation. the upper boundary is placed at the base of the lowest highly fossiliferous sandstone bed of the muslingeelv member. distribution. same as the formation. chronostratigraphy. ryazanian, praetollia maynci and hectoroceras kochi ammonite zones. muslingeelv member history. erected by surlyk et al. (1973) and described in more detail by surlyk (1973). type section. north-east slope of the mountain of crinoid bjerg, immediately west of muslingeelv 70°29.7′n, 23°04.0′w (surlyk et al. 1973, fig. 36a; figs 2a, 73a). reference section. north-east slope of the mount of crinoid bjerg (surlyk et al. 1973, fig. 36b) (fig. 73b). thickness. 10−35 m. lithology. the basal unit of fossiliferous, commonly pebbly sandstones is overlain by massive or cross-bedded sandstones. in the western part of the canyon fill, the fossiliferous beds are succeeded by thin mudstones followed by massive or large-scale cross-bedded sandstones. fossils. ammonites, bivalves, belemnites, wood, trace fossils. depositional environment. shelf-margin canyon fill. boundaries. the lower boundary is placed at the base of the lowest fossiliferous pebbly sandstone bed, overlying silty and sandy mudstones of the crinoid bjerg member or locally the raukelv formation. the member forms the uppermost unit in the area, so the nature of the original upper boundary is not known. distribution. same as the formation. chronostratigraphy. ryazanian, hectoroceras kochi ammonite zone. hall bredning group new group history. erected here to include the mainly upper jurassic mudstone-dominated part of the jurassic succession on the east coast of milne land, representing the milne land subbasin. name. after hall bredning, that part of the scoresby sund fjord, which borders eastern milne land (fig. 1). type area. eastern milne land (figs 1, 2b). thickness. up to about 1200 m. lithology. the lower part of the group is dominated by conglomerates and coarse-grained sandstones, onlapping a south-eastward inclined crystalline basement surface. the middle part is dominated by black mudstones and dark grey silty mudstones with intercalated sandstones in some members. the upper part marks a return to fineto coarse-grained sandstones. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 83 of 116 www.geusbul let in.org fossils. ammonites, belemnites, corals, the bivalve buchia, and rich bivalve faunas in some coarser-grained units, dinoflagellate cysts, plant fragments and trace fossils. depositional environment. marine rocky shoreline at the base, followed by middle and lower shoreface, offshore transition zone and offshore, returning to shoreface and shelf-margin wedge, and topped by shoreface and paralic sediments. boundaries. unconformably overlies crystalline basement and forms the top stratum in the outcrop area in eastern milne land. unconformably overlain by palaeogene plateau basalts in the western part of the outcrop area. distribution. eastern milne land (fig. 1). possibly also south-east gåseland (håkansson et al. 1971). subdivision. the group is subdivided into the charcot bugt, kap leslie, hartz fjeld and pinnadal formations. chronostratigraphy. lower bathonian – hauterivian, based on rich ammonite faunas and dinoflagellate cysts. key references. aldinger (1935), spath (1935, 1936), håkansson et al. (1971), callomon & birkelund (1980, 1982), fürsich (1982), fürsich & heinberg (1983), birkelund et al. (1984), birkelund & callomon (1985), larsen (1995), surlyk (2003). charcot bugt formation history. the formation was erected by aldinger (1935) and revised and described by callomon & birkelund (1980). type section. visdal, south-east milne land. sections m44 (70°39.4′n, 25°46.7′w) and m49 (70°41.1′n, 25°49.0′w) of birkelund et al. (1984) are here designated as a combined type section (figs 1, 2b, 75–77). reference sections. parat kløft, kosmocerasdal and mudderbugt (birkelund et al. 1984, sections m1, m2 and m39, respectively; fig. 2b). thickness. up to 200 m. lithology. conglomerates with kaolinised pebbles at the base, overlain by several sets of cross-bedded and largescale high-angle clinoform-bedded sandstones that are pebbly in places. fossils. rare ammonites, bivalves, corals, belemnites, plant fragments and trace fossils. si f m c pb co cl s 0 20 40 60 80 100 120 m ka p le sli e fm (p ar s) c ha rc ot b ug t f m vi sd al m b fig 75 fig. 75 composite type section of the charcot bugt formation and the visdal member in the visdal valley, milne land (figs 1, 2b). for legend, see fig. 7. 0 2 4 6 c ha rc ot b ug t f m (p ar s) 8 f cl si s m c pb m fig 76 fig. 76 detailed log of the basal beds of the charcot bugt formation (and the visdal member) forming part of the composite type section of these units in the visdal valley, milne land (figs 1, 2b). from larsen et al. (2003, fig. 6). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 84 of 116 www.geusbul let in.org depositional environment. transgressive shoreface off a rocky coast and prograding shelf-margin wedges. boundaries. the lower boundary is a major unconformity and is placed where conglomerates and coarsegrained sandstones onlap an irregular, locally kaolinised inclined surface of crystalline basement rocks (fig. 76). the upper boundary is placed at the base of dark grey laminated mudstones of the kap leslie formation, overlying coarse-grained, locally pebbly sandstones; the boundary is a sharp diachronous drowning surface. distribution. south-eastern milne land and possibly also south-eastern gåseland (fig. 1, 2b). scree covered 60 40 30 20 10 0 50 m 110 100 90 80 m 4 4 7 4 8 11 11 7 8 11 8 9 9 7 4 4 4 8 7 8 8 4 5 5 4 7 120 cl si f m cgr sand cl si f m cgr sand kap leslie fm c ha rc ot b ug t f m vi sd al m b (p ar s) c ha rc ot b ug t f m vi sd al m b (p ar s) fig 77 fig. 77 section through the upper part of the visdal member, charcot bugt formation forming part of the composite type section of these units, corresponding to locality m44 of birkelund et al. (1984). visdal, milne land (figs 1, 2b). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 85 of 116 www.geusbul let in.org chronostratigraphy. lower bathonian – middle oxfordian, the a. arcticus, a. ishmae, a. cranocephaloide and a. densiplicatum ammonite zones, faunal horizons m-1 and m-7 of callomon & birkelund (1980) and j-9, j-14 and j-18 of callomon (1993); dinoflagellate cyst assembages equivalent to the q. mariae and c. cordatum ammonite zones in the upper part (larsen et al. 2003). total span is thus the a. arcticus to c. densiplicatum ammonite zones. subdivision. the formation is subdivided (from below) into the visdal and mudderbugt members. key references. aldinger (1935), håkansson et al. (1971), beauvais (1977), callomon & birkelund (1980), larsen (1995), larsen et al. (2003). visdal member history. erected by callomon & birkelund (1980). equivalent to the entire charcot bugt formation with the exception of the uppermost 6 m exposed at southern visdal, which are referred to the mudderbugt member. type section. visdal, south-east milne land. sections m44 (70°39.4′n, 25°46.7′w) and m49 (70°41.1′n, 25°49.0′w) of birkelund et al. (1984), designated as the combined type section of the charcot bugt formation, are similarly defined as the type section of the visdal member (figs 1, 2b, 75–79). reference section. due to the westward onlap of the formation, the lowest part of the member is probably located to the east in parat kløft (birkelund et al. 1984, section m1) although poorly exposed (fig. 2b). thickness. up to 200 m. lithology. as the formation. characteristic yellowish white and red weathering colours. fossils. as in the formation. a distinctive conglomeratic bed composed of basement clasts, rounded corals, oyster shells and rare ammonites was described from the northern end of visdal by callomon & birkelund (1980). depositional environment. as for the formation. boundaries. the lower boundary is a distinct onlap surface where weathered crystalline basement is overlain by coarse-grained, locally pebbly sandstones and conglomerates of the visdal member (fig. 76). the upper boundary is placed at the base of the coarseto very coarse-grained pebbly sandstones of the mudderbugt member in the southern end of the visdal valley and elsewhere at the base of dark mudstones of the kap leslie formation. the upper boundary is strongly diachronous, younging to the west. the member interfingers with the kap leslie formation to the east, and in sections of alternating sandstone–mudstone units the boundary is placed at the 50% cut-off position, above which point mudstones dominate. distribution. same as the formation. chronostratigraphy. lower bathonian – middle oxfordian, a. arcticus – c. densiplicatum ammonite zones; faunal horizon m-1 of callomon & birkelund (1980) has been identified in the lower part of the member based on ammonites. dinoflagellate cysts recovered from the uppermost clinoform set in visdal indicate the middle oxfordian, corresponding to faunal horizon m-7 of callomon & birkelund (1980). fig. 78 sandstones of the visdal member, charcot bugt formation, viewed looking north. this section corresponds to locality m44 of birkelund et al. (1984), forming an important part of the composite type section of the charcot bugt formation and the visdal member. overlain with a sharp boundary by dark mudstones of the kosmocerasdal member. visdal, milne land (figs 1, 2b). http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 86 of 116 www.geusbul let in.org mudderbugt member history. erected by callomon & birkelund (1980). type section. southern visdal, milne land, 70°38.5′n, 25°48.0′w (callomon & birkelund 1980, section 40; figs 1, 2b, 80). thickness. up to 6 m. lithology. light grey weathering, very coarse-grained, pebbly trough cross-bedded sandstones, capped by a pebble lag. fossils. bivalves, rare ammonites and trace fossils. depositional environment. shoreface. boundaries. the lower boundary is placed at the base of the lowermost light grey weathering, trough cross-bedded, very coarse-grained and pebbly sandstone, overlying the yellow weathering visdal member. the upper boundary is placed at the sharp base of dark mudstones of the kap leslie formation, overlying a pebble lag; this boundary represents a marked flooding surface. fig 79_version 2 fig. 79 upper part of the visdal member at the type locality, comprising a 40 m thick clinoform bed, overlain at a sharp boundary by dark mudstones of the kosmocerasdal member (kap leslie formation). corresponds to locality m44 of birkelund et al. (1984, see figs 77, 78). visdal, milne land (figs 1, 2b). si f m c pbcl s 0 10 20 m ka p le sli e fm c ha rc ot b ug t f m m ud de rb ug t m b vi sd al m b (p ar s) ko sm oc er as da l m b (p ar s) fig 80 fig. 80 type section of the mudderbugt member (charcot bugt formation), southern visdal, milne land (figs 1, 2b). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 87 of 116 www.geusbul let in.org distribution. southern end of the visdal valley, milne land (fig. 2b). chronostratigraphy. middle oxfordian, c. densiplica­ tum ammonite zone (fauna horizon m-7 of callomon & birkelund 1980), based on ammonites. kap leslie formation history. erected by rosenkrantz (1929), revised and described by callomon & birkelund (1980) and birkelund et al. (1984). type area. the eastern slope of hartz fjeld is here designated as the type area (see birkelund et al. 1984, fig. 3; fig. 2b). for representative sections, see under the individual members. not all members have measured logs due to the fine-grained nature of the succession, often leading to poor exposure. the formation has a long history of study and the style of the published logs is thus variable. thickness. up to 800 m, thins to 70–80 m west of visdal. lithology. dark grey laminated mudstones, grey bioturbated sandy siltstones, subordinate fineto med iumgrained sandstones, commonly glauconitic. for details, see the component members. fossils. abundant ammonites, belemnites, bivalves, spores, pollen, dinoflagellate cysts and trace fossils; vertebrates, crustaceans, brachiopods, gastropods, serpulids and wood occur locally. depositional environment. lower shoreface – offshore. boundaries. the lower boundary is placed at the base of dark grey mudstones, overlying a pebble lag that caps sandstones of the charcot bugt formation, representing a marked drowning surface. the upper boundary is placed at the base of the lowest sandstone bed of the hartz fjeld formation, overlying dark grey mudstones of the astartedal member. distribution. eastern milne land (figs 1, 2b). chronostratigraphy. middle callovian – middle volgian, s. calloviense – c. anguinus ammonite zones, faunal horizon m-35 of callomon (1993) and faunal horizons m-2 to m-46 of birkelund et al. (1984), based on ammonites and dinoflagellate cysts. subdivision. the formation is subdivided (from below) into the kosmocerasdal, aldinger elv, bays elv, cardioceraskløft, gråkløft, krebsedal, pernaryggen, and astartedal members. key references. aldinger (1935), spath (1935, 1936), callomon & birkelund (1980, 1982), piasecki (1979), birkelund et al. (1984), birkelund & callomon (1985), fürsich (1982), fürsich & heinberg (1983), larsen et al. (2003). kosmocerasdal member history. erected by callomon & birkelund (1980). type section. kosmocerasdal (70°44.6′n, 25°29.1′w; fig. 2b) and the coastal section 1 km south of nordøstelv (70°45.3′n, 25°18.3′w) are here designated a combined type section (birkelund et al. 1984, sections m2 and m4, respectively; fig. 81). thickness. up to 170 m but wedges out westward. lithology. grey bioturbated sandy siltstones, with subordinate thin sandstone beds that increase in abundance in the upper levels of the member. fossils. ammonites, belemnites, bivalves and dinoflagellate cysts. depositional environment. lower shoreface – offshore transition zone. boundaries. the lower boundary coincides with the lower boundary of the formation and is placed at the sharp base of dark mudstones overlying sandstones of the charcot bugt formation, at a major flooding surface. the upper levels of the member show a progressive increase in interbedded sandstones, grading into the poorly consolidated sands of the aldinger elv member. the boundary is placed where the proportion of sandstone to siltstone exceeds 50%. distribution. eastern milne land, most thickly developed in the charcot bugt area (figs 1, 2b). chronostratigraphy. middle callovian – upper oxfordian, s. calloviense – a. glosense ammonite zones, faunal horizon j-35 of callomon (1993) and faunal horizons m-2 to m-10 of callomon & birkelund (1980) and birkelund et al. (1984), based on ammonites and indirectly on dinoflagellate cysts. lower boundary diachronous, younging westwards to the upper oxfordian. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 88 of 116 www.geusbul let in.org aldinger elv member history. erected by callomon & birkelund (1980) for the pecten sandstone of aldinger (1935) and described by fürsich & heinberg (1983). type section. a composite type section is designated here (fig. 82), combining a section in cardioceraskløft, 70°44.8′n, 25°18.6′w (fig. 2b; fürsich & heinberg 1983, fig. 1, section 5) and a section between cardioceraskløft csmsfs ssicl 0 20 40 60 80 100 120 140 180 160 m ka p le sli e fm charcot bugt fm (pars) ko sm oc er as da l m b al di ng er e lv m b fig 81 fig. 81 type section of the kosmocerasdal member (kap leslie formation), kosmocerasdal, milne land (figs 1, 2b). modified from callomon & birkelund (1980, fig. 2); re-measured by s. piasecki. for legend, see fig. 7. vf f m s c 0 10 20 30 40 50 60 kosmocerasdal mb (pars) bays elv mb (pars) al di ng er e lv m b ka p le sli e fm 70 80 90 m fig 82 fig. 82 type section of the aldinger elv member (kap leslie formation). milne land (figs1, 2b). schematic composite section. modified from fürsich & heinberg (1983, fig. 3). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 89 of 116 www.geusbul let in.org and aldinger elv, 70°42.1′n, 25°30.8′w (fig. 2b; fürsich & heinberg 1983, fig. 1, section 25; birkelund et al. 1984, fig. 2, section m25). thickness. thins from about 70 m in the east to a few metres in the west. lithology. fineto medium-grained, poorly cemented cross-bedded sandstones (fig. 83). fossils. abundant bivalves and serpulids, echinoderms, ammonites, trace fossils and wood fragments. depositional environment. interpreted as an offshore bar by fürsich & heinberg (1983), and as a shelf-margin wedge by surlyk (2003). boundaries. the lower boundary is placed approximately at the 50% cut-off position between sandy siltstones and well-bedded sandstones of the kosmocerasdal member and the loose sandstones of the aldinger elv member. the upper boundary is sharp and placed at the base of alternating sandy siltstones and fine-grained sandstones of the bays elv member (fig. 83). distribution. eastern milne land (figs 1, 2b). chronostratigraphy. upper oxfordian, upper a. glosense – a. serratum ammonite zones, faunal horizon m-11 of callomon & birkelund (1980) and birkelund et al. (1984), based on ammonites and indirectly on dinoflagellate cysts. bays elv member history. erected by callomon & birkelund (1980). type section. cardioceraskløft, eastern milne land, 70°44.2′n, 25°19.6′w, is here designated the type section (birkelund et al. 1984, section m6; figs 1, 2b, 84). thickness. up to 20 m. lithology. alternating sandy siltstones and fine-grained glauconitic sandstones. fossils. few ammonites and bivalves, dinoflagellate cysts. depositional environment. offshore transition zone. boundaries. the lower boundary is sharp and placed at the base of sandy siltstones, overlying sandstones of the aldinger elv member. the upper boundary is placed at the top of a widely recognisable green glauconitic sandstone bed, overlain by grey, bioturbated siltstones. distribution. eastern milne land (figs 1, 2b). chronostratigraphy. upper oxfordian – lowermost kimmeridgian, a. regulare – p. baylei zones, faunal horizons m-12 to m-14 of callomon & birkelund (1980) and birkelund et al. (1984), based on ammonites and indirectly on dinoflagellate cysts. fig. 83 yellow sandstones of the uppermost aldinger elv member (c. 6 m exposed), sharply overlain by black mudstones of the bays elv member (both kap leslie formation). hartz fjeld, milne land (figs 1, 2b). http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 90 of 116 www.geusbul let in.org cardioceraskløft member history. erected by callomon & birkelund (1980). type section. cardioceraskløft, eastern milne land, 70°44.2′n, 25°19.6′w, is here designated the type section (birkelund et al. 1984, section m6; figs 1, 2b, 84). thickness. 30 m. lithology. grey, highly micaceous, bioturbated siltstones with abundant horizons of calcareous concretions. fossils. ammonites, gastropods and dinoflagellate cysts. depositional environment. offshore marine. boundaries. the lower boundary is placed at the base of the lowest grey, bioturbated siltstone, overlying alternating sandy siltstones and fine-grained, glauconitic sandstones of the bays elv member. the upper boundary is placed where black laminated mudstones, referred to the gråkløft member, overlie grey bioturbated siltstones of the cardioceraskløft member. distribution. eastern milne land (figs 1, 2b). chronostratigraphy. lower kimmeridgian, r. cymodoce – a. mutabilis ammonite zones, faunal horizons m-15 to m-19 of birkelund et al. (1984), based on ammonites and indirectly on dinoflagellate cysts. gråkløft member history. erected by callomon & birkelund (1980). type section. northern slope of gråkløft, eastern milne land, 70°43.8′n, 25°19.0′w (birkelund et al. 1984, fig. 4, section m8; figs 1, 2b, 85). thickness. 160 m in the east, wedging out towards the west. lithology. black, laminated, organic-rich mudstone with pyritic beds and concretions. fossils. ammonites, bivalves at a few levels. depositional environment. offshore marine; deposited under dysoxic to anoxic conditions. boundaries. the lower boundary is placed at the base of the lowermost laminated black mudstones, overlying grey, bioturbated siltstones. the upper boundary is placed at the base of the lowermost very fine-grained muddy sandstones referred to the krebsedal member. distribution. eastern milne land (figs 1, 2b). chronostratigraphy. kimmeridgian – lower volgian, a. eudoxus – p. elegans ammonite zones, faunal horizons m-20 to m-24 of birkelund et al. (1984), based on ammonites and indirectly on dinoflagellate cysts. ba ys e lv m b al di ng er e lv m b (p ar s) c ar di oc er as kl øf t m b ka p le sli e fm g rå kl øf t m b (p ar s) si f m c cl s fig 84 0 10 20 30 40 50 m fig. 84 type sections of the bays elv and cardioceraskløft members (both kap leslie formation) at cardioceraskløft, milne land (figs 1, 2b). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 91 of 116 www.geusbul let in.org krebsedal member history. erected by birkelund et al. (1984). type section. ridge south of krebsedal, eastern milne land, 70°41.9′n, 25°18.4′w (birkelund et al. 1984, section m17; figs 1, 2b, 86). thickness. 160 m at the type section, ranges between 100 and 200 m. lithology. grey, very fineto fine-grained muddy sandstone. quartz pebbles and glaucony grains recorded locally in the basal levels. the upper part of the member at the kronen mountain consists of large-scale cross-bedded sandstone (birkelund et al. 1984). fossils. ammonites, belemnites, bivalves, crustaceans, ichthyosaurs, plesiosaurs, dinoflagellate cysts and trace fossils. depositional environment. marine, offshore transition zone to lower shoreface. boundaries. the lower boundary is placed where finegrained muddy sandstones overlie black laminated mudstones of the gråkløft member. the upper boundary is placed at the base of the lowermost glauconitic sandstone of the pernaryggen member; this boundary is diachronous. distribution. eastern milne land (figs 1, 2b). 0 10 20 30 40 m m m m 50 60 70 80 90 100 110 120 130 140 150 160 fig 85 si s si s si s si s c ar di oc er as kl øf t m b (p ar s) g rå kl øf t m b ka p le sli e fm g rå kl øf t m b ka p le sli e fm g rå kl øf t m b ka p le sli e fm g rå kl øf t m b kr eb se da l m b (p ar s) ka p le sli e fm fig. 85 type section of the gråkløft member (kap leslie formation) at hartz fjeld, milne land (figs 1, 2b). modified from birkelund et al. (1984, fig. 4). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 92 of 116 www.geusbul let in.org chronostratigraphy. lower–middle volgian, p. wheatley­ ensis – d. liostracus ammonite zones, faunal horizons m-25 to m-37 of birkelund et al. (1984). pernaryggen member history. erected by birkelund et al. (1984). type section. pernaryggen, eastern milne land, 70°43.0′n, 25°26.6′w (birkelund et al. 1984, section m23; figs 1, 2b, 87). thickness. 53 m at the type section, thins westward to 15 m at bay fjelde. lithology. bioturbated, micaceous and glauconitic sandstone dominate the member (fig. 88). large-scale cross-bedded sandstones are prominent at the base of the type section. fossils. ammonites, belemnites, bivalves, brachiopods, serpulids, gastropods, crustaceans and trace fossils. depositional environment. lower and upper shoreface. boundaries. the lower boundary is placed at the lowest glauconitic sandstone, overlying finer-grained nonglauc onitic sandstones of the krebsedal member. the fig 86 sicl kr eb se da l m b ka p le sli e fm g rå kl øf t m b (p ar s) pe rn ar yg ge n m b (p ar s) s 0 50 100 150 m fig. 86 type section of the krebsedal member (kap leslie formation) in the krebsedal valley, milne land (figs 1, 2b). modified from birkelund et al. (1984, fig. 7). for legend, see fig. 7. m si s kr eb se da l m b (p ar s) pe rn ar yg ge n m b h en ni gr yg ge n m b (p ar s) ka p le sli e fm h ar tz f je ld f m 0 10 20 30 40 50 fig 87 fig. 87 type section of the pernaryggen member (kap leslie formation) at pernaryggen, milne land (figs 1, 2b). modified from birkelund et al. (1984, fig. 8). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 93 of 116 www.geusbul let in.org upper boundary is placed at the base of the lowest silty micaceous mudstone of the astartedal member, if present, or at the erosional unconformity beneath the yellow sandstones of the hartz fjeld formation. distribution. milne land (figs 1, 2b). chronostratigraphy. middle volgian, p. rugosa – c. angui­ nus ammonite zones, faunal horizons m-33 to m-45 of birkelund et al. (1984). astartedal member history. erected by birkelund et al. (1984). type section. ridge south of krebsedal, eastern milne land, 70°41.9′n, 25°18.4′w (birkelund et al. 1984, section m17; figs 1, 2b, 89). thickness. 20–26 m in the type area, wedges out towards the north. lithology. light grey, micaceous, silty very fine-grained bioturbated loose sands. large-scale cross-bedding is locally recognisable. fossils. ammonites, belemnites, bivalves, trace fossils, plant debris and wood fragments. depositional environment. shoreface and offshore transition zone. boundaries. the lower boundary is placed at the base of the lowest light grey, bioturbated, very fine-grained sandstones, overlying glauconitic sandstones of the pernaryggen member; at most localities, this is a sharp boundary. the upper boundary is an erosional fig 88 pi kr he pe fig. 88 ridge exposing micaceous and glauconitic sandstones of the pernaryggen member (pe) of the kap leslie formation. photograph taken from the mountain of kronen towards hartz fjeld, where the hennigryggen (he) and kronen (kr) members of the hartz fjeld formation are exposed, capped by the pinnadal formation (pi). milne land (figs 1, 2b). http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 94 of 116 www.geusbul let in.org unconformity at the base of yellow non-micaceous quartz sandstones of the hartz fjeld formation. distribution. very restricted extent in the hartz fjeld area, eastern milne land (figs 1, 2b). chronostratigraphy. middle volgian. l. groenlandicus ammonite zone, faunal horizon m-46 of birkelund et al. (1984). hartz fjeld formation history. erected by rosenkrantz (1929), described by sykes & brand (1976), and subdivided by birkelund et al. (1984) into the hennigryggen and pinnadal members. the hennigryggen member of birkelund et al. (1984) is redefined herein, and the pinnedal member is elevated to formation status. type section. the type section is defined by its constituent members. thickness. about 300 m. lithology. fineto coarse-grained sandstone showing cross-bedding and high-angle clinoform bedding. fossils. bivalves, rare ammonites, brachiopods, plant fragments and trace fossils. depositional environment. shoreface, shelf-margin wedge. boundaries. the lower boundary is placed at the base of cross-bedded or clinoform-bedded sandstones, unconformably overlying micaceous, silty mudstones or loose sands of the kap leslie formation. the formation forms the uppermost stratal units on the kronen and bay fjelde mountains. it is succeeded by the pinnadal formation on the mountain of hartz fjeld where the boundary is placed at the base of a prominent dark grey silty mudstone unit, overlying pebbly sandstones or a marked pebble lag. distribution. milne land (figs 1, 2b). chronostratigraphy. middle volgian – valanginian, based on ammonites and dinoflagellates. subdivision. the formation is subdivided into the hennigryggen member and the overlying new kronen member. an important upper volgian – ryazanian hiatus separates the lower and the upper parts of the hennigryggen member as originally defined by birkelund et al. (1984). the hennigryggen member is here redefined to only include the lower part, whereas the upper part is referred to the new kronen member. the hennigryggen member is thus of middle volgian age, whereas the kronen member is of valanginian age. the pinnadal member of birkelund et al. (1984) is herein excluded from the hartz fjeld formation and is elevated to the rank of formation. key references. sykes & brand (1976), piasecki (1979), birkelund et al. (1984), surlyk & noe-nygaard (1995), surlyk (2003). hennigryggen member revised member history. erected by birkelund et al. (1984). the hennigryggen member as defined by birkelund et al. (1984) conceals a major hiatus, corresponding to an important si f m c co cl s pb as ta rt ed al m b pe rn ar yg ge n m b (p ar s) h en ni gr yg ge n m b (p ar s) 0 10 20 30 m ka p le sli e fm h ar tz f je ld f m fig 89 fig. 89 type section of the astartedal member (kap leslie formation) in the krebsedal valley, milne land (figs 1, 2b). modified from birkelund et al. (1984, fig. 10). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 95 of 116 www.geusbul let in.org unconformity covering all of the upper volgian – ryazanian (birkelund et al. 1984). the member is thus redefined here, being restricted to only include the lower part, below the unconformity. type section. north-east corner of the mountain of hartz fjeld, eastern milne land, 70°43.2′n, 25°20.4′w (birkelund et al. 1984, section m9, equivalent to the ‘upper section’ of sykes & brand 1976; figs 1, 2b, 90). si pbs si pbs si pbs as ta rt ed al m b h en ni gr yg ge n m b kr on en m b kr on en m b pi nn ad al fm h ar tz f je ld f m h ar tz f je ld f m ka p le sli e fm h ar tz f je ld f m 0 10 20 30 40 50 60 70 80 90 100 120 140 160 180 200 220 240 260 280 300 m m m fig 90 fig. 90 type sections of the hennigryggen and kronen members, both hartz fjeld formation. hartz fjeld, milne land (figs 1, 2b). modified from birkelund et al. (1984, fig. 11). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 96 of 116 www.geusbul let in.org fig. 91 light grey, silty, very fine-grained sands of the pernaryggen member (kap leslie formation, persons for scale standing on this member) overlain by yellow sandstones of the hennigryggen member, hartz fjeld formation. the mountain of kronen (west of hartz fjeld), milne land, looking towards the west (figs 1, 2b). fig. 92 photograph of the thick-bedded hennigryggen member. the sandstone-dominated succession (up to the dark ledges) corresponds to the lowest c. 70 m of the succession shown in fig. 90. western side of the hartz fjeld mountain, milne land (figs 1, 2b). http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 97 of 116 www.geusbul let in.org thickness. about 200 m. lithology. stacked coarsening-upward, cross-bedded and high-angle clinoform-bedded, commonly glauconitic sandstones (figs 91, 92). fossils. bivalves, rare ammonites, brachiopods, plant fragments, and trace fossils. depositional environment. marine shoreface and shelf margin wedge. boundaries. the lower boundary coincides with that of the formation. the upper boundary is placed at the base of cross-bedded sandstones of the new kronen member at an unconformity that represents a major hiatus. distribution. the member is known from the mountains of hartz fjeld, kronen and bay fjelde in milne land (figs 1, 2b). chronostratigraphy. middle volgian, l. groenlandicus ammonite zone, faunal horizon m-47 of birkelund et al. (1984). kronen member new member history. this new member formed the upper part of the hennigryggen member of previous usage (birkelund et al. 1984). a major hiatus corresponding to all of the upper volgian and the ryazanian occurs in the middle of the member as originally defined, and the hennigryggen member is revised here to only include the part below the unconformity. the upper part is here defined as the new kronen member. type section. same locality as for the hennigryggen member. north-east corner of the mountain of hartz fjeld, eastern milne land, 70°43.2′n, 25°20.4′w (birkelund et al. 1984, section m9, equivalent to the ‘upper section’ of sykes & brand 1976; figs 1, 2b, 90). reference section. the mountain of kronen, west of hartz fjeld, 70°42.9′n, 25°27.6′w (birkelund et al. 1984, section m23; figs 1, 2b). thickness. about 130 m based on our interpretation of data in birkelund et al. (1984). lithology. stacked coarsening-upward sandstone-dominated units showing trough and low-angle cross-bedding and thick clinoform beds; some coarsening-upward units have a basal siltstone interval. in many cases, these units are capped by a ferruginous crust that may be associated with pebble beds; the trace fossil tisoa habichi extends down from the top of many such units. carbonaceous matter occurs scattered. post-depositional decalcification has resulted in the loss of shelly fossils and the obliteration of sedimentary structures in many beds. fossils. ammonites, trace fossils including tisoa habichi, planolites, curvolithus and gyrochorte, casts of wood logs. depositional environment. progradational marine shorefaces and shelf-margin wedges. boundaries. the lower boundary is placed at the top of a marked pebble lag that caps the upper coarsening-upwards sandstone unit of the redefined hennigryggen member; this lag is overlain by cross-bedded sandstones referred to the kronen member. the upper boundary is placed where coarsening-upwards sandstones are succeeded by alternating silty mudstones and coarse-grained sheet sandstones of the lower pinna dal formation. distribution: the hartz fjeld and kronen mountains, milne land (figs 1, 2b). chronostratigraphy. lowermost valanginian tollia kli­ movskiensis ammonite zone has been identified (faunal horizon m-48 of callomon & birkelund 1982). the top of the redefined hennigryggen member contains middle volgian ammonites, whereas the basal beds of the new kronen member contain early valanginian or possibly late ryazanian ammonites. the hiatus between the two members thus comprises all of the upper volgian substage and all or most of the ryazanian stage. an important hiatus may also be present between the kronen member and the overlying pinnadal formation, as much of the valanginian is apparently absent. biostratigraphical data from the pinnadal formation are, however, very sparse. pinnadal formation revised unit history. erected as a member by birkelund et al. (1984). elevated herein to the rank of formation. type locality. pinnadal, north-east corner of hartz fjeld, 70°43.1′n 25°20.4′w (sykes & brand 1976, fig. 1; figs 1, 2b). the locality is poorly known; representative detailed http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 98 of 116 www.geusbul let in.org sections modified from sykes & brand (1976, fig. 6) are shown in fig. 93. thickness. 102 m at the type section. lithology. units of interbedded micaceous, silty mudstones and sandstones alternate with coarsening-upwards units of coarse-grained sandstones and occasional conglomerates. thin coal beds are recorded. fossils. trace fossils and dinoflagellate cysts. depositional environment. paralic, marginal marine to freshwater, possibly lagoonal, as indicated by palynomorphs, alternating with possible fan deltas (sykes & brand 1976; piasecki 1979). boundaries. the lower boundary is placed at the base of a prominent dark grey silty mudstone bed, overlying pebbly sandstones or a marked pebble lag capping the kronen member. the formation forms the top stratum of the sedimentary succession in milne land. distribution. hartz fjeld, milne land (figs 1, 2b). chronostratigraphy. hauterivian, on the basis of dinoflagellate cysts (piasecki 1979). wollaston forland group history. erected by surlyk (1978a). type area. wollaston forland (fig. 1). thickness. estimated to be up to 3 km. lithology. dominated by breccias, conglomerates, pebbly sandstones and dark sandy and silty mudstones of the lindemans bugt formation, overlain by thinner conglomerates and sandstones, calcareous light grey and red mudstones of the palnatokes bjerg formation. the lithologies of the group and its constituent formations and members are described in detail by surlyk (1978a, 1984), surlyk & korstgård (2013), henstra et al. (2016) and hovikoski et al. (2018). fossils. ammonites, belemnites, bivalves and trace fossils at some levels. depositional environment. marine fault-scarp talus, coalescent submarine fans forming slope aprons, basinfloor fans, and condensed mudstones over submerged half-graben crests. boundaries. the lower boundary is placed at the base of the lowest breccias, conglomerates and pebbly sandstones, overlying yellow sandstones of the pelion formation, dark mudstones of the bernbjerg formation, 0 5 10 m 0 5 m 0 5 m cl si s cl si s cl si s fig 93 pi nn ad al fm (p ar s) pi nn ad al fm (p ar s) pi nn ad al fm (p ar s) fig. 93 representative sections from the type locality of the pinnadal formation (hartz fjeld formation) in the pinnadal valley, milne land (figs 1, 2b). modified from sykes & brand (1976, fig. 6). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 99 of 116 www.geusbul let in.org or crystalline basement with marked erosional and in some cases angular unconformity. the upper boundary may be unconformable or conformable, dependent on the structural setting, and is placed at the base of the lower cretaceous mudstones of the brorson halvø group of bjerager et al. (2020). distribution. wollaston forland, kuhn ø, th. thomsen land, hochstetter forland and store koldewey. on traill ø, the group is only represented by the valanginian albrechts bugt and rødryggen members of the palnatokes bjerg formation (figs 1, 2d, 2e). chronostratigraphy. middle volgian – hauterivian, based on ammonites and bivalves. subdivision. the group is subdivided (from below) into the lindemans bugt and palnatokes bjerg formations. key references. vischer (1943), maync (1947, 1949), donovan (1953, 1964), surlyk & clemmensen (1975), surlyk (1975, 1978a, 1978c, 1984, 1989, 2003), nøhr-hansen (1993), kelly et al. (1998), alsen (2006), alsen & mutterlose (2009), pauly et al. (2012a, b, 2013), surlyk & korstgård (2013), möller et al. (2015), henstra et al. (2016), hovikoski et al. (2018), bjerager et al. (2020). lindemans bugt formation history. erected by surlyk (1978a). type area. north-west wollaston forland (fig. 1); see component members for the type sections. reference sections. niesen (surlyk 1978a, sections 15, 20, 22, 24, 34), palnatoke bjerg (surlyk 1978a, sections 49, 50; fig. 2e). thickness. up to about 2 km. lithology. breccias, boulder, cobble and pebble conglomerates, pebbly sandstones, coarse-grained sandstones, dark silty or sandy mudstones. fossils. ammonites, belemnites, the bivalve buchia, plant fragments. depositional environment. marine fault-scarp talus, slope apron, coalescent submarine fans, basin plain. boundaries. the lower boundary is the same as for the group. the upper boundary is placed at the base of the lowermost light grey mudstones and thin finegrained sandstones of the palnatokes bjerg formation, fig 94a_ny a fig 94b_ny b fig. 94 type section of the laugeites ravine member (lindemans bugt formation), laugeites ravine, kuhn ø (figs 1, 2e). the upper part of the member is shown in (a), the middle part in (b), in both exposures, the exposed section is about 10 m thick. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 100 of 116 www.geusbul let in.org overlying the conglomerate-dominated lindemans bugt formation. distribution. wollaston forland, kuhn ø, th. thomsen land and store koldewey (figs 2d, 2e). chronostratigraphy. middle volgian – middle upper ryazanian, based on ammonites and bivalves. subdivision. the formation is subdivided (from below) into the laugeites ravine, rigi and niesen members. laugeites ravine member history. erected by surlyk (1978a). type section. laugeites ravine, southern kuhn ø (figs 1, 2e, 94). reference sections. laugeites ravine (surlyk 1978a, sections 5, 7; 74°47.9′n, 20°33.5′w), niesen (surlyk 1978a, section 201; figs 2e, 95). thickness. at least 80 m, possibly much more. lithology. interlaminated dark silty–sandy mudstones and thin sandstones, pebbly sandstones and pebble conglomerates. fossils. ammonites, belemnites, bivalves and plant fragments. depositional environment. base-of-slope, fan fringe and interfan slope. boundaries. the lower boundary is placed where the lowest bed of interlaminated silty–sandy mudstones and thin sandstones, pebbly sandstones or pebble conglomerates overlie black, laminated, better sorted and finer grained mudstones of the bernbjerg formation. the upper boundary is placed at the base of the lowermost cobble or boulder conglomerate bed referred to the rigi member. distribution. wollaston forland and kuhn ø (figs 1, 2e). chronostratigraphy. middle–upper volgian, based on ammonites and bivalves. rigi member history. erected by surlyk (1978a). type area. north-west slopes of the palnatoke bjerg and niesen mountains, wollaston forland (figs 1, 2e). fig 95 0 si s 5 10 15 20 25 m li nd em an s b ug t f m la ug ei te s r av in e m b (p ar s) fig. 95 reference section of the laugeites ravine member (lindemans bugt formation), laugeites ravine, kuhn ø (figs 1, 2e). from surlyk (1978a, section 7). for legend, see fig. 7. 0 fig 96 li nd em an ns b ug t f m ri gi m b (p ar s) 5 10 15 20 25 30 m cl si s pbco b fig. 96 reference section from the type area of the thickly developed rigi member (lindemans bugt formation) at the niesen mountain, wollaston forland (figs 1, 2e). from surlyk (1978a, section 22). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 101 of 116 www.geusbul let in.org reference sections. niesen, (surlyk 1978a, sections 15, 22), palnatoke bjerg (surlyk 1978a, sections 30–34, 49, 50; figs 2e, 96, 97). thickness. estimated to be about 2 km. lithology. breccias, conglomerates, pebbly and coarsegrained sandstones. fossils. ammonites, belemnites, bivalves and plant fragments. depositional environment. fault-scarp talus, slope apron, coalescent submarine fans and basin-floor fans. boundaries. the lower boundary is placed where the member unconformably overlies crystalline basement rocks, where the lowest cobble or boulder conglomerate bed overlies interbedded dark mudstones and 0 10 15 20 25 30 35 40 45 50 m 5 55 65 70 75 80 85 90 95 100 105 m 60 fig 97 cl si s pbcob cl si s pbcob li nd em an s b ug t f m ri gi m b (p ar s) li nd em an s b ug t f m ri gi m b (p ar s) fig. 97 reference section from the type area of the thickly developed rigi member (lindemans bugt formation), palnatoke bjerg, wollaston forland (figs 1, 2e). from surlyk (1978a, section 34). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 102 of 116 www.geusbul let in.org sandstones of the laugeites ravine member, or where the base of the conglomerate-dominated member rests unconformably on yellow sandstones of the pelion formation or black mudstones of the bernbjerg formation. the upper boundary coincides with the upper boundary of the formation. distribution. wollaston forland, th. thomsen land, kuhn ø (figs 1, 2e). chronostratigraphy. middle volgian – upper ryazanian, based on ammonites and bivalves. niesen member history. erected by surlyk (1978a). type section. niesen, wollaston forland, 74°39.4′n, 20°31.2′w (surlyk 1978a, section 203; figs 1, 2e, 98). reference section. store koldewey (for details, see bjerager et al. 2020; fig. 1). thickness. up to 200 m, wedges out towards the east. lithology. dark mudstones, interlaminated mudstones and fine-grained sandstones. fossils. ammonites, belemnites, bivalves and plant fragments. depositional environment. base-of-slope and basin plain. boundaries. the lower boundary is placed at the base of dark mudstones and other fine-grained lithologies, overlying the conglomerate-dominated rigi member. the upper boundary is placed where dark mudstones are succeeded by the conglomeratic young sund member or by light grey mudstones of the albrechts bugt member, both belonging to the upper ryazanian – hauterivian palnatokes bjerg formation. distribution. wollaston forland, store koldewey (figs 1, 2d, 2e). chronostratigraphy. on wollaston forland, upper volgian – upper ryazanian, based on ammonites and bivalves. lower–upper ryazanian on store koldewey based on ammonites and dinoflagellates. palnatokes bjerg formation history. erected by surlyk (1978a). type section. niesen, wollaston forland, 74°38.3′n, 20°29.8′w (figs 1, 2e, 99). reference sections. eastern kuhn ø (surlyk 1978b, sections 10, 13), niesen (surlyk 1978a, section 18), rødryggen (surlyk 1978a, sections 53, 54), store koldewey (bjerager et al. 2020; figs 1, 2e). thickness. about 600 m on wollaston forland, wedging out to the east to 25–30 m. up to 70 m on store koldewey. lithology. conglomerates, pebbly sandstones, sandstones, interlaminated mudstones and sandstones, dark mudstones, light grey mudstones with abundant calcareous concretions and concretionary layers, yellow and dark red mudstones. fossils. bivalves and belemnites locally very common, ammonites relatively rare. m 50 30 20 10 0 fig 98 cl si s li nd em an s b ug t f m n ie se n m b (p ar s) fig. 98 type section of the niesen member (lindemans bugt formation), niesen, wollaston forland (figs 1, 2e). from surlyk (1978a, section 203). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 103 of 116 www.geusbul let in.org 0 4 8 12 16 20 24 26 28 clay silt vf f m c vc sand fig 99 30 32 38 42 46 50 54 58 m clay silt vf f m c vc sand clay silt vf f m c vc sand 120 126 130 134 138 142 146 150 m 58 62 66 70 74 78 82 86 90 94 98 102 106 110 114 118 120 a.b. yo un g su nd m b yo un g su nd m b (p ar s) a. b. /y .s. pa ln at ok es b je rg f m pa ln at ok es b je rg f m pa ln at ok es b je rg f m al br ec ht s b ug t m b (p ar s) al br ec ht s b ug t m b yo un g su nd m b fig. 99 type section of the palnatokes bjerg formation and reference section of the young sund member of this formation at the niesen mountain, wollaston forland (figs 1, 2e). a.b.: thin tongues of the albrechts bugt member. a.b./y.s.: transitional young sund and albrechts bugt members. from hovikoski et al. (2018, fig. 6). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 104 of 116 www.geusbul let in.org depositional environment. slope apron, basin-floor fan, basin plain, submerged intrabasinal highs formed over half-graben crests. boundaries. on wollaston forland, the lower boundary is placed where thinly interbedded light grey mudstones and fine-grained sandstones overlie the conglomerate dominated lindemans bugt formation. where the formation rests on dark mudstones of the niesen member of the lindemans bugt formation, the boundary is placed at the base of mainly light grey mudstones with more concretionary horizons. on store koldewey, the lower boundary is placed at the base of a varicoloured greyish–reddish conglomerate bed where it overlies crystalline basement, yellow sandstones of the pelion formation, or grey muddy sandstones of the lindemans bugt formation. the palnatokes bjerg formation is overlain by dark barremian–albian mudstones of the stratumbjerg formation of bjerager et al. (2020). this boundary is typically an erosional unconformity but may be locally conformable on wollaston forland and store koldewey (bjerager et al. 2020). distribution. wollaston forland and kuhn ø. isolated outliers on traill ø, hochstetter forland and store koldewey (figs 1, 2d, 2e). chronostratigraphy. on wollaston forland, uppermost ryazanian (bojarkia mesezhnikovi ammonite zone) – hauterivian (simbirskites beds), based on ammonites and bivalves. on store koldewey, uppermost ryazanian (bojarkia mesezhnikovi ammonite zone) – lowermost upper barremian based on ammonites, belemnites, bivalves, nannofossils and dinoflagellates. subdivision. the formation is subdivided into the young sund, falskebugt, albrechts bugt and rødryggen members in the wollaston forland and kuhn ø areas, and into the midter gneisnæs and ravn pynt members on store koldewey. young sund member history. erected by surlyk (1978a). type section. stratumbjerg on the southern side of the cardiocerasdal valley, south-west wollaston forland, 74°26.6′n, 20°13.7′w (surlyk 1978a, section 37; figs 1, 2e, 100). reference section. niesen, wollaston forland (figs 2e, 99). thickness. the original maximum thickness is estimated at about 600 m. less than 50 m at the type locality. interfingers with the albrechts bugt member through a thickness of about 250–300 m in northern wollaston forland (fig. 99), and about 150 m on store koldewey. lithology. conglomerates with subordinate breccias, coarseand medium-grained sandstones with thin pebble beds. the clasts in the conglomerates are normally well-rounded quartzites, gneisses and amphibolites. in the lower levels of the member, the conglomerates and breccias contain sandstone and heterolith blocks or rafts derived from the pelion and jakobsstigen formations; clasts of light grey mudstone and calcareous concretions, both from the albrechts bugt member, are also common. fossils. bivalves and belemnites. cl si pb cos 40 35 30 25 20pa ln at ok es b je rg f m yo un g su nd m b (p ar s) jak ob ss tig en f m (p ar s) 15 10 5 0 m fig 100 fig. 100 type section of the young sund member at stratumbjerg, southern side of cardiocerasdal, wollaston forland (figs 1, 2e). from surlyk (1978a, section 37). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 105 of 116 www.geusbul let in.org depositional environment. slope apron and basin-floor fan. boundaries. in south-west wollaston forland, the lower boundary is placed at an angular unconformity at the base of the lowest conglomerates overlying yellow sandstones of the pelion formation, mudstones of the bernbjerg formation, or sandstones and heteroliths of the jakobsstigen formation (figs 100, 101). where it overlies the rigi member of the lindemans bugt formation, it can be distinguished by its large content of reworked clasts of sedimentary rocks, notably light grey mudstones and concretions of the penecontemporaneous albrechts bugt member. the upper boundary is unconformable and placed at the base of dark barremian–albian mudstones with intercalated sandstones of the stratumbjerg formation of bjerager et al. (2020). the member interfingers laterally to the east with the albrechts bugt member and in transitional sections, the boundary is placed at the 50% cut-off position between conglomerates characterising the young sund member and mudstones or fine-grained sandstones typical of the albrechts bugt member. distribution. western wollaston forland, clavering ø (figs 1, 2e). chronostratigraphy. uppermost ryazanian – valanginian, based on ammonites and bivalves. falskebugt member history. erected by surlyk (1978a). type locality. western slope of falkebjerg, north-eastern wollaston forland, 74°34.9′n. 19°20.9′w (figs 1, 2e, 102). thickness. about 40–50 m. lithology. conglomerates and coarse-grained sandstones, which may contain reworked albrechts bugt member sediments (fig. 102). fossils. bivalves, brachiopods. boundaries. rests on caledonian crystalline basement rocks and interfingers to the west with albrechts bugt member mudstones. in transitional sections, the boundary between the two members is placed at the 50% cut-off position between the conglomerates and sandstones of the falskebugt member and the mudstones of the albrechts bugt member. the upper boundary is unconformable and is placed at the base of the dark barremian–albian mudstones of the stratumbjerg formation of bjerager et al. (2020). distribution. north-east wollaston forland (figs 1, 2e). chronostratigraphy. valanginian–hauterivian(?), based on bivalves and ammonites. piasecki et al. (2020) have recently suggested that the uppermost levels of this member may interfinger with the stratumbjerg formation, though this interpretation is based primarily on a poorly exposed area. if confirmed, however, this relationship would indicate a barremian age for the uppermost falskebugt member. young sund mb sst. w e cgl. cgl. jakobsstigen fm (jurassic) fig 101 fig. 101 submarine gravity-flow conglomerates of the young sund member (palnatokes bjerg formation) at the type section, erosionally overlying upper jurassic sandstones of the jakobsstigen formation. stratumbjerg, south-western wollaston forland (figs 1, 2e), figure (right) for scale. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 106 of 116 www.geusbul let in.org albrechts bugt member history. erected by surlyk (1978a). type section. rødryggen, north-central wollaston forland, 74°32.7′n, 19°50.9′w (figs 1, 2e, 103, 104). reference sections. perisphinctes ravine, wollaston forland (surlyk 1978a, section 10), eastern kuhn ø (surlyk 1978a, section 13), niesen, wollaston forland (surlyk 1978a, section 18), rødryggen, wollaston forland (surlyk 1978a, sections 53, 54; figs 2e, 99). thickness. about 300 m in northern wollaston forland where it interfingers with the most fine-grained parts of the young sund member. thins to the east to about 30 m. lithology. sandy, light-grey to yellowish mudstones with abundant calcareous concretions, nodules and layers. fossils. ammonites, bivalves, belemnites. brachiopods, calcareous nannofossils and trace fossils. depositional environment. basin floor, hanging-wall dip slope and submerged intrabasinal highs formed over fault-block crests. boundaries. the lower boundary is placed at the base of light grey mudstones, overlying darker niesen member mudstones, which contain much fewer concretions, or conglomerates and sandstones of the rigi member. in southern and eastern wollaston forland and eastern kuhn ø, the boundary is placed at the base of light grey mudstones, overlying black mudstones of the bernbjerg formation with angular or erosional unconformity. the member interfingers laterally with conglomerates and sandstones of the young sund member to the west and with conglomerates and sandstones of the falskebugt member to the east; in such situations, the member boundaries are defined at the 50% cut-off point between the respective lithologies. the upper boundary is placed at the base of red mudstones of the rødryggen member and is sharp or transitional. in the latter case, the base of the lowest red mudstone defines the boundary. the upper boundary is unconformable in some areas and placed at the base of dark mudstones of the stratumbjerg formation of bjerager et al. (2020). distribution. western, central and south-west wollaston forland and probable outliers south of falskebugt, on eastern kuhn ø and eastern hochstetter forland. small outliers occur on north-east traill ø (figs 1, 2e). chronostratigraphy. uppermost ryazanian – hauterivian, based on ammonites and bivalves. rødryggen member history. erected by surlyk (1978a). type section. rødryggen, north-central wollaston forland, 74°32.7′n, 19°50.9′w (surlyk 1978a, section 54; figs 1, 2e, 103, 104). fig 102 fig. 102 type locality of the falskebugt member (lindemans bugt formation), hammer (encircled) for scale. no section measured due to lack of reliable continuous exposure. western slope of the falkebjerg mountain, north-eastern wollaston forland (figs 1, 2e). photograph by jørgen bojesen-koefoed. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 107 of 116 www.geusbul let in.org reference sections. rødryggen, central wollaston forland (surlyk 1978a, section 53), and stratumbjerg, south-west wollaston forland (figs 1, 2e). thickness. up to about 13 m. lithology. red mudstones with intercalated fine sandy yellow mudstones (figs 104, 105). fossils. bivalves, brachiopods, ammonites, belemnites, echinoids, crinoids and trace fossils. depositional environment. submerged intrabasinal highs formed over fault block crests. boundaries. the lower boundary with the underlying albrechts bugt member is placed at the base of the lowermost red mudstone bed. in some areas, the lower boundary is placed at an unconformity on black mudstones of the bernbjerg formation. the upper boundary is placed at the base of the unconformably overlying dark mudstones of the stratumbjerg formation of bjerager et al. (2020). 0 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 m 16 17 18 19 21 22 23 24rø dr yg ge n m b al br ec ht s b ug t m b pa ln at ok es b je rg f m pa ln at ok es b je rg f m al br ec ht s b ug t m b 25 26 27 m be rn bj er g fm . 20 cl si f m c gr scl si f m c gr s fig 103 fig. 103 type sections of the albrechts bugt and rødryggen members at rødryggen, wollaston forland (figs 1, 2e). from alsen (2006, fig. 28 and appendix 1.4). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 108 of 116 www.geusbul let in.org distribution. stratumbjerg in south-west wollaston forland, central wollaston forland, eastern kuhn ø, small outliers in eastern hochstetter forland and north-east traill ø (figs 1, 2e). chronostratigraphy. upper ryazanian – upper valanginian on traill ø, hauterivian in the wollaston forland area, based on ammonites and bivalves. fig 104 be ab st rø fig. 104 oblique aerial photograph of the rødryggen ridge, wollaston forland (figs 1, 2e), viewed towards the south-east, showing mudstones of the bernbjerg formation (be) overlain by the palnatokes bjerg formation comprising light-coloured calcareous mudstones of the albrechts bugt member (ab), followed by red mudstones of the rødryggen member (rø, c. 10 thick). the palnatokes bjerg formation is succeeded by dark mudstones of the stratumbjerg formation (st). dark mudstones in the badlands in the background belong to the bernbjerg formation. fig 105 ab rø st fig. 105 the type locality for the albrechts bugt member (ab) and the rødryggen member (rø), both palnatokes bjerg formation. succeeded by dark mudstones of the stratumbjerg formation (st). rødryggen, wollaston forland (figs 1, 2e). http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 109 of 116 www.geusbul let in.org midter gneisnæs member history. erected by bjerager et al. (2020) as a new member of the palnatokes bjerg formation to encompass very coarse siliciclastic sediments in the store koldewey area that were earlier assigned to the ‘aucella conglomerate’ and the cape hamburg formation by koch (1929). the latter name was abandoned by surlyk (1978a), due to lack of stratigraphic information. type section. trækpasset, store koldewey, 76°09.6′n, 18°34.4′w, where the geometry of the member is well exposed (figs 2d, 106, 107). reference section. midter gneisnæs, store koldewey (figs 1, 2d). thickness. at least 30 m at midter gneisnæs and 48 m at trækpasset on store koldewey, within the axes of depositional channels or gullies, thinning out to less than 1 m at channel margins. up to a few metres thick elsewhere on store koldewey. lithology. the thicker, axial successions consist of coarse conglomerates and breccias, few fossil-rich pebble conglomerates (fig. 107), coarseand medium-grained sandstones with thin pebble beds and subordinate mudstone beds. at ravn ravine, and at channel margins, the member is about one metre thick and consists of muddy and sandy conglomerates with abundant fossils, predominantly buchia bivalves. fossils. bivalves, belemnites and ammonites are common in certain intervals. depositional environment. submarine gully/channel, and channel margins, submarine fault-scarp slope apron and fan. boundaries. the lower boundary is placed at the base of a varicoloured greyish–reddish conglomerate, overlying crystalline basement or yellow sandstones of the pelion formation with an erosional contact. the upper boundary is placed at the base of grey mudstones of the ravn pynt member, which in places has a sandy basal part, or the member is unconformably overlain by yellowish grey sandstones of the basal stratumbjerg formation (bjerager et al. 2020). distribution. store koldewey (figs 1, 2d). chronostratigraphy. the age of the lower part is not well constrained. the middle and upper parts of the member are assigned to the uppermost ryazanian – upper valanginian based on ammonites, belemnites, bivalves, palynomorphs and nannoplankton. fig 106 cl si s gn pb vf f m cvc f m c f 0 10 20 30 40 50 60 70 m pa ln at ok es b je rg f m m id te r g ne isn æs m b (p ar s) ra vn p yn t m b (p ar s) fig. 106 type section of the midter gneisnæs member at trækpasset, store koldewey (figs 1, 2d). from bjerager et al. (2020). for legend, see fig. 7. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 110 of 116 www.geusbul let in.org ravn pynt member history. erected by bjerager et al. (2020) for the upper part of the palnatokes bjerg formation on store koldewey (fig. 1). type section. ravn pynt, store koldewey, 76°08.9′n, 18°33.2′w (bjerager et al. 2020; fig. 2d). reference section. ravn ravine, store koldewey (figs 1, 2d). thickness. 10–100 m thick. lithology. bioturbated grey silty and sandy mudstones with thin sandstone beds. fossils. large inoceramid bivalves and belemnites common, rare ammonites, common wood fragments and dinoflagellates. trace fossils include zoophycos, nereites, diplocraterion, palaeophycus and rare paleodictyon. depositional environment. deep-water slope gully and slope. boundaries. the lower boundary is poorly exposed at the type section but is well exposed in the reference section at ravn ravine, where it is placed at the base of a grey mudstone bed, overlying a fossiliferous conglomerate of the midter gneisnæs member (palnatokes bjerg formation). in the type section, the upper boundary is placed at the erosional base of a pebbly and fossil-rich fig. 107 midter gneisnæs member at the type section, showing stacked conglomerate – pebbly sandstone beds. trækpasset, store koldewey (figs 1, 2d). http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 111 of 116 www.geusbul let in.org sandstone bed of the basal stratumbjerg formation. in the reference section at ravn ravine, the upper boundary is placed at the base of dark grey laminated mudstones of the stratumbjerg formation (bjerager et al. 2020). distribution. store koldewey (figs 1, 2d). chronostratigraphy. lower valanginian – lower barremian, based on ammonites, belemnites, bivalves, dinoflagellates and nannoplankton. acknowledgements the comprehensive stratigraphic revision presented here results from the work of the senior author and a number of colleagues and graduate students since the late 1980s. we are grateful for generous economic support from the danish natural science research council, the carlsberg foundation, norsk hydro, statoil, saga petroleum, the norwegian petroleum directorate, amoco, conoco and bp. we acknowledge the late j.h. callomon for assistance in determination of ammonites from critical stratigraphic levels. we are very grateful to jette halskov for artwork. we thank the reviewers snorre olaussen and harald brekke for penetrating and painstaking, yet constructive comments. special thanks go to geus scientific editor j.r. ineson, not only for comments on an early manuscript version but also for his dedicated editorial work on the final manuscript. additional information funding economic support from the danish natural science research council, the carlsberg foundation, norsk hydro, statoil, saga petroleum, the norwegian petroleum directorate, amoco, conoco and bp. author contributions fs planned, structured and wrote the integrated manuscript, in the latter stages with the assistance of pa. all authors, to varying degrees, contributed with primary data and initial texts to the description of individual lithostratigraphic units, and reviewed relevant parts of the final manuscript prior to submission. http://www.geusbulletin.org surlyk et al. 2021: geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 112 of 116 www.geusbul let in.org ahokas, j.m., nystuen, j.p. & martinius, a.w. 2014: stratigraphic signatures of punctuated rise in relative sea-level in an estuary-dominated heterolithic succession: incised valley fills of the toarcian ostreaelv formation, neill klinter group (jameson land, east greenland). marine and petroleum geology 50, 103–129. https://doi.org/10.1016/j. marpetgeo.2013.11.001 aldinger, h. 1935: geologische beobachtungen im oberen jura des scoresbysundes (ostgrönland). meddelelser om grønland 99, 128 pp. alsen, p. 2006. the early cretaceous (late ryazanian – early hauterivian) ammonite fauna of north-east greenland: taxonomy, biostratigraphy, and biogeography. fossils and strata 53, 229 pp. alsen, p. & mutterlose, j. 2009: the early cretaceous of north-east greenland: a crossroads of belemnite migration. palaeogeography, palaeoclimatology, palaeoecology 280, 168–182. https://doi. org/10.1016/j.palaeo.2009.06.011 alsen, p. & piasecki, s. 2018: biostratigraphy of the hareelv formation (upper jurassic) in the blokelv-1 core, jameson land, central east greenland. in: ineson, j. & bojesen-koefoed, j.a. 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(eds): the jurassic of north-east greenland. geological survey of denmark and greenland bulletin 5, 9–18. https://doi.org/10.34194/ geusb.v5.4800 williford, k.h., grice, k., holman, a. & mcelwain, j.c. 2014: an organic record of terrestrial ecosystem collapse and recovery at the triassic– jurassic boundary in east greenland. geochimica et cosmochimica acta 127, 251–263. https://doi.org/10.1016/j.gca.2013.11.033 http://www.geusbulletin.org jurassic stratigraphy of east greenland abstract introduction geological setting revised lithostratigraphy jameson land supergroup kap stewart group innakajik formation primulaelv formation rhætelv formation neill klinter group rævekløft formation gule horn formation elis bjerg member albuen member ostreaelv formation horsedal member astartekløft member lepidopteriselv member nathorst fjeld member harris fjeld member skævdal member trefjord bjerg member sortehat formation vardekløft group bristol elv formation bastians dal formation muslingebjerg formation pelion formation ugleelv member parnas member spath plateau member fossilbjerget formation goniomyakløft member payer dal formation jakobsstigen formation olympen formation athene member hades member zeus member hareelv formation bernbjerg formation ugpik ravine member raukelv formation sjællandselv member salix dal member langelandselv member prasino member fynselv member kokkino member straight river member rauk plateau member hesteelv formation crinoid bjerg member muslingeelv member hall bredning group charcot bugt formation visdal member mudderbugt member kap leslie formation kosmocerasdal member aldinger elv member bays elv member cardioceraskløft member gråkløft member krebsedal member pernaryggen member astartedal member hartz fjeld formation hennigryggen member kronen member pinnadal formation wollaston forland group lindemans bugt formation laugeites ravine member rigi member niesen member palnatokes bjerg formation young sund member falskebugt member albrechts bugt member rødryggen member midter gneisnæs member ravn pynt member acknowledgements additional information references figures fig. 1: simplified geological map of the east greenland rift basin. see the figure caption for full fig. 2: maps showing the position of place names and localities mentioned in the text. see the figur fig. 2: maps showing the position of place names and localities mentioned in the text. see the figur fig. 2: maps showing the position of place names and localities mentioned in the text. see the figur fig. 3: stratigraphical scheme. see the figure caption for full details. fig. 3: stratigraphical scheme. see the figure caption for full details. fig. 4: chronostratigraphic scheme. see the figure caption for full details. fig. 5: lateral profile of the type section of the innakajik formation. see the figure caption for f fig. 6: photo of conglomerates, pebbly sandstones and compressed coalified branches. see the figure fig. 7: type section of the primulaelv formation. see the figure caption for full details. fig. 7: type section of the primulaelv formation. see the figure caption for full details. fig. 7: type section of the primulaelv formation. see the figure caption for full details. fig. 8: photo showing the primulaelv formation at astartekløft. see the figure caption for full deta fig. 9: photo showing the boundary between the kap stewart and neill klinter groups. see the figure fig. 10: type section of the rhætelv formation. see the figure caption for full details. fig. 11: photo of the rhætelv formation. see the figure caption for full details. fig. 12: aerial photograph of the neill klinter group at harris fjeld. see the figure caption for fu fig. 12: aerial photograph of the neill klinter group at harris fjeld. see the figure caption for fu fig. 13: reference section of the rævekløft formation. see the figure caption for full details. fig. 14: reference section of the gule horn formation and type section of the ostreaelv formation. s fig. 15: type section of the elis bjerg member. see the figure caption for full details. fig. 16: type section of the albuen member. see the figure caption for full details. fig. 17: type section of the horsedal member. see the figure caption for full details. fig. 18: type section of the astartekløft member. see the figure caption for full details. fig. 19: type section of the lepidopteris elv member. see the figure caption for full details. fig. 20: type section of the nathorst fjeld member. see the figure caption for full details. fig. 21: type section of the harris fjeld member. see the figure caption for full details. fig. 22: type section of the skævdal member. see the figure caption for full details. fig. 23: type section of the trefjord bjerg member. see the figure caption for full details. fig. 24: photo showing boundary between the skævdal member and the trefjord bjerg member. see the fi fig. 25: photo showing the boundary between the trefjord bjerg member and the sortehat formation. se fig. 26: log of a cored section of the sortehat formation at the type locality. see the figure capti fig. 27: reference section of the sortehat formation. see the figure caption for full details. fig. 28: type section of the bristol elv formation. see the figure caption for full details. fig. 29: photo showing type section of the bristol elv formation. see the figure caption for full de fig. 30: type section of the bastians dal formation. see the figure caption for full details. fig. 31: photo showing fluvial conglomerates and sandstonesof the bastians dal, formation. see the f fig. 32: type section of the muslingebjerg formation. see the figure caption for full details. fig. 33: photo showing the type section of the muslingebjerg formation. see the figure caption for f fig. 34: type section of the pelion formation. see the figure caption for full details. fig. 35: photo showing the type section of the pelion formation. see the figure caption for full det fig. 36: type section of the ugleelv member. see the figure caption for full details. fig. 37: photo showing the north slope of the ugleelv valley, southern jameson land. see the figure fig. 38: type section of the parnas member. see the figure caption for full details. fig. 39: type section of the spath plateau member. see the figure caption for full details. fig. 40: aerial photograph of the type section of the spath plateau member. see the figure caption f fig. 41: type section of the fossilbjerget formation. see the figure caption for full details. fig. 42: type section of the goniomyakløft member. see the figure caption for full details. fig. 43: type section of the payer dal formation. see the figure caption for full details. fig. 44: field photograph of the type section of the payer dal formation. see the figure caption for fig. 45 type section of the jakobsstigen formation. see the figure caption for full details. fig. 46: aerial photograph of the type section of the jakobsstigen formation. see the figure caption fig. 47: type section of the jakobsstigen formation. see the figure caption for full details. fig. 48: type section of the olympen formation and its constituent athene, hades and zeus members. s fig. 49: photo showing the sandstones of the athene membersee the figure caption for full details. fig. 50: type section of the hareelv formation. see the figure caption for full details. fig. 51: reference section of the hareelv formation. see the figure caption for full details. fig. 52: photo showing the hareelv formation. see the figure caption for full details. fig. 53: type section of the bernbjerg formation. see the figure caption for full details. fig. 54: reference section of the bernbjerg formation. see the figure caption for full details. fig. 55: type section of the ugpik ravine member. see the figure caption for full details. fig. 56: raukelv formation stratigraphic relationships. see the figure caption for full details. fig. 57: type sections of the sjællandselv and salix dal members, raukelv formation, salix dal. see fig. 58: photo showing the boundary of the well-structured sandstones of the fynselv member. see the fig. 59: type sections of the langelandselv and prasino members and reference section of the fynselv fig. 60: reference sections of the langelandselv and prasino members. see the figure caption for ful fig. 61: reference sections of the prasino and fynselv members. see the figure caption for full deta fig. 62: line drawings from field photographs of the prasino member. see the figure caption for full fig. 63: photo showing large-scale cross-bedded lower part of the prasino member. see the figure cap fig. 64: close-up of section shown in fig. 63. see the figure caption for full details. fig. 65: reference section of the fynselv member. see the figure caption for full details. fig. 66: photo showing the plateau-forming high-angle clinoform bed of the fynselv member. see the f fig. 67: photo showing the fynselv member. see the figure caption for full details. fig. 68: type section of the kokkino member. see the figure caption for full details. fig. 69: type section of the straight river member. see the figure caption for full details. fig. 70: reference section of the straight river member. see the figure caption for full details. fig. 71: photograph of the type section of the rauk plateau member. see the figure caption for full fig. 72: photograph of the high-angle clinoform-bedded rauk plateau member. see the figure caption f fig. 73: hesteelv formation. see the figure caption for full details. fig. 74: contoured block diagram of a major erosional canyon incised in the se margin of the prograd fig. 75: composite type section of the charcot bugt formation and the visdal member. see the figure fig. 76: detailed log of the basal beds of the charcot bugt formation. see the figure caption for fu fig. 77: section through the upper part of the visdal member. see the figure caption for full detail fig. 78: sandstones of the visdal member. see the figure caption for full details. fig. 79: upper part of the visdal member at the type locality. see the figure caption for full detai fig. 80: type section of the mudderbugt member. see the figure caption for full details. fig. 81: type section of the kosmocerasdal member. see the figure caption for full details. fig. 82: type section of the aldinger elv member. see the figure caption for full details. fig. 83: yellow sandstones of the uppermost aldinger elv member. see the figure caption for full det fig. 84: type sections of the bays elv and cardioceraskløft members. see the figure caption for full fig. 85: type section of the gråkløft member. see the figure caption for full details. fig. 86: type section of the krebsedal member. see the figure caption for full details. fig. 87: type section of the pernaryggen member. see the figure caption for full details. fig. 88: photo showing a ridge exposing micaceous and glauconitic sandstones of the pernaryggen memb fig. 89: type section of the astartedal member. see the figure caption for full details. fig. 90: type sections of the hennigryggen and kronen members. see the figure caption for full detai fig. 91: photo showing light grey, silty, very fine-grained sands of the pernaryggen member. see the fig. 92: photograph of the thick-bedded hennigryggen member. see the figure caption for full details fig. 93: representative sections from the type locality of the pinnadal formation. see the figure ca fig. 94: type section of the laugeites ravine member. see the figure caption for full details. fig. 95: reference section of the laugeites ravine member. see the figure caption for full details. fig. 96: reference section from the type area of the rigi member. see the figure caption for full de fig. 97: reference section from the type area of the rigi member. see the figure caption for full de fig. 98: type section of the niesen member. see the figure caption for full details. fig. 99: type section of the palnatokes bjerg formation and reference section of the young sund memb fig. 100: type section of the young sund member. see the figure caption for full details. fig. 101: submarine gravity-flow conglomerates of the young sund member. see the figure caption for fig. 102: type locality of the falskebugt member. see the figure caption for full details. fig. 103: type sections of the albrechts bugt and rødryggen members at rødryggen. see the figure cap fig. 104: oblique aerial photograph of the rødryggen ridge. see the figure caption for full details. fig. 105: the type locality for the albrechts bugt member and the rødryggen member. see the figure c fig. 106: type section of the midter gneisnæs member. see the figure caption for full details. fig. 107: photo showing the midter gneisnæs member at the type section. see the figure caption for f korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 1 of 96 monograph the kangâmiut dykes in west greenland: markers of the tectono-metamorphic evolution of the southern nagssugtoqidian orogen and its foreland john a. korstgård1 , flemming c. mengel*2 , william e. glassley†3, kai sørensen4 1department of geoscience, aarhus university, denmark; 21795 view point rd, boulder, co, 80305, usa; 3department of earth and planetary sciences, university of california, davis, ca, 95616, usa, and center for earth system petrology, department of geoscience, aarhus university, denmark († deceased); 4brønshøjholms allé 53, dk-2700 brønshøj, denmark abstract the general extent and structural evolution of the southern nagssugtoqidian orogen of west greenland were first described by hans ramberg who based much of his paper on the deformation of the regional kangâmiut dyke swarm. the southern boundary is marked by a transition from undeformed, discordant dykes in the south to highly deformed dykes and host rocks to the north. our analysis of the southern nagssugtoqidian orogen and its southern foreland uses a comprehensive compilation of available data and covers the area from sisimiut in the north to alanngua, south of maniitsoq. this represents almost the entire c. 200 km latitudinal extent of the kangâmiut dyke swarm and encompasses the complete range of nagssugtoqidian overprint on these dykes and their country rocks. south of itillip ilua (itilleq), the structural and metamorphic overprints on the dykes exhibit a considerable range in both intensity and p–t conditions between and even within outcrops. in contrast, north of itillip ilua, the rocks show more systematic gradual increases in the degree of structural overprints and metamorphic grade, culminating in the ikertooq thrust zone where granulite facies rocks are brought southwards over amphibolite facies rocks. currently, available age data from the nagssugtoqidian orogen permits the identification of two metamorphic episodes at c. 1850–1800 ma and c. 1780–1720 ma. these groups of metamorphic ages are supported by recent 40ar–39ar ages from dykes in the same area, which cluster at c. 1860 ma and c. 1740 ma, respectively. albeit geographically sporadic, both age intervals support a subdivision of the nagssugtoqidian structural and metamorphic overprints across the southern nagssugtoqidian orogen and its foreland into two distinguishable temporal phases. further geochronological investigations may well, however, find these two phases to be part of a tectonic continuum. for now, it is thought that the older event records south-directed thrusting over the foreland and concomitant loading of this crust, at least as far south as maniitsoq. this c. 1860–1800 ma crustal shortening and thrusting likely also closed a depositional basin located at the current latitude of ikertooq, which could have formed during an early-orogenic extensional event that enabled and accompanied the c. 2035 ma emplacement of kangâmiut dykes. up to 50–100 ma later, a younger (c. 1780–1720 ma) phase of shearing and thrusting mainly affected the itillip ilua – ikertooq area and likely overprinted elements of the former event. this local younger overprint generated a separate trend of distinctly northward-increasing deformation and metamorphism. *correspondence: campmengel@sbcglobal.net received: 22 feb 2022 revised: 08 jul 2022 accepted: 19 jul 2023 published: 31 may 2024 keywords: paleoproterozoic tectonics, southern nagssugtoqidian orogen, kangâmiut dykes, south-directed thrusting, west greenland abbreviations: agp: amphibole-garnet geobarometer apt: amphibole-plagioclase geothermometer cno: central nagssugtoqidian orogen e-morb: enriched mid-ocean ridge basalts hree: heavy ree icp-ms: inductively coupled plasma-mass spectroscopy kbar: kilobar lass: laser ablation split stream lree: light ree lil: large-ion lithophile loi: loss on ignition n-morb: normal mid-ocean ridge basalts nno: northern nagssugtoqidian orogen p-t: pressure–temperature ree: rare-earth element shrimp: sensitive high resolution ion microprobe sno: southern nagssugtoqidian orogen tims: thermal ionization mass spectrometry tiat: ti-in-amphibole geothermometer t-morb: transitional mid-ocean ridge basalts geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: thomas find kokfelt (geus, denmark), kerstin saalmann (ngu, norway) reviewed by: martin klausen (stellenbosch university, south africa) and one anonymous reviewer. funding: none. competing interests: see page 90 additional files: see page 90 https://doi.org/10.34194/geusb.v58.8312 https://orcid.org/0000-0003-2143-8514 https://orcid.org/0000-0001-5028-9998 https://orcid.org/0009-0009-8938-092x mailto:campmengel@sbcglobal.net korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 2 of 96 geusbulletin.org 1 introduction basic dyke swarms occur in many precambrian terranes and act as sensitive indicators of the post-emplacement tectonic and metamorphic overprints affecting the rocks that the dykes intruded. this is due to the generally large areal extent of the dykes and their overall planar structural nature, as well as the limited primary chemical variations among the members of the swarm. the global co-occurrence of c. 2400– 2000 ma dyke swarms and 1900–1700 ma orogenic activity across many precambrian terranes are generally interpreted to indicate stages of initial break-up of precambrian continents (or supercontinents), and the subsequent re-assembly of these continental masses through orogenic convergence and collision (e.g. ernst et al. 1995). unraveling the complex evolution of these orogenic belts has often benefited from detailed studies of the structural and metamorphic effects recorded in mafic dyke swarms. examples include the characterisation of laxfordian orogeny by studying the structural and metamorphic overprints on the scourie dykes in the lewisian gneiss complex of nw scotland (e.g. sutton & watson 1951; heaman & tarney 1989). similarly, analyses of orogenic overprints on the kikkertavak dykes, the avayalik dykes and the domes dykes in the margins of the nain province of labrador, canada (e.g. ryan et al. 1983; korstgård & ermanovics 1984; ryan 1990; ermanovics & ryan 1990; mengel et al. 1991) were instrumental in outlining the nature and architecture of the paleoproterozoic torngat and makkovik orogens of the area. within this framework, the aim of this study is twofold: first, to assess the extent and nature of nagssugtoqidian metamorphic overprints across the southern nagssugtoqidian orogen (sno) and its southern archaean foreland by studying the kangâmiut dykes of west greenland (fig. 1). second, to use that information in combination with structural, metamorphic and geochronological data to provide constraints on the dynamic evolution of this area. in this contribution, we compile previously published studies of the kangâmiut dyke swarm in west greenland (65°10’n to 67°n), and present an extensive database including previously unpublished research, which allows us to refine our understanding of the tectonic and metamorphic evolution of the southern part of the nagssugtoqidian orogen and its foreland. after summarising the regional geology of the nagssugtoqidian orogen and its southern foreland (chapter 2), we describe in detail the kangâmiut dykes and their host rocks between maniitsoq in the south of the study area and itillip ilua (itilleq) to the north (chapters 3, 4), as well as the more intensely deformed and metamorphosed dykes and gneisses between itillip ilua and ikertooq (chapters 5, 6). the geochemical and metamorphic features of the kangâmiut dykes are described in chapter 7, thus providing a background for the metamorphic pressure and temperature calculations (chapter 8). lastly, we discuss and summarise the results and implications of this study. an overview of the place names referred to throughout this manuscript is provided in fig. 2. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 3 of 96 geusbulletin.org 50 km ma�c dykes supracrustal rocks variably reworked archean gneisses kangâmiut ma�c dyke swarm sisimiut charnockite a arfersior�k quartz diorite thrust fault undi�erentiated archean/paleoproterozoic paleoproterozoic s sisimiut sisimiut isortuat kangaamiut sermiat nf greenland ice sheet attu qeqertarsuup tunua qasigiannguit ilulissat aasiaat itilleq kangaamiut maniitsoq so ut he rn a rc ha ea n fo re la nd nordre strøm�ord shear zone nordre isorto q shear zone ikertooq shear zone s c n o n n o sn o a tuno terrane a kia terrane alanngua c omplex iso rto q qorlortoq gneiss kangerlussuatsiaq ikertooq kangerlu ssu aq nassuttooq itillip ilua 52° 67° 68°68° 67° 52° 50° 69° 66° 54° fig. 1 simplified geological map of west greenland. solid black lines indicate the internal subdivisions of the nagssugtoqidian orogen, including the sno (southern nagssugtoqidian orogen), cno (central nagssugtoqidian orogen) and nno (northern nagssugtoqidian orogen; after marker et al. 1995). black dashed lines separate the akia terrane, alanngua complex and tuno terrane (locations from steenfelt et al. 2021) and correspond to the previously named nordland, alângua and kangâmiut complexes, respectively, of noe-nygaard & ramberg (1961a). solid green lines indicate the kangâmiut dykes investigated in this study from coastal regions (65°10’n to 67°n). north of itillip ilua and the nagssugtoqidian front (nf), country rocks are largely archaean gneisses and minor supracrustals, all variably to completely reworked during nagssugtoqidian orogenesis. the trace of the nf is from hageskov (1995, 1997, 1998) and olsen (2000). solid grey line: aujassoq–evighedsfjord shear zone (allaart & jensen 1979). grey dashed lines: nordre strømfjord shear zone and nordre isortoq shear zone. map based on information from escher (1971), allaart (1982) and marker et al. (1995). the extent of the nagssugtoqidian orogen across southern greenland is shown in the inset map (grey shading). https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 4 of 96 geusbulletin.org n f fig. 9 fig. 12 fig. 15 fig. 16 fig. 35 fig. 38 sarfartoq inussuttusup tunua ungusivik kangerlu ssu aq sulussugut itilliup qaqqaa akuliaruseq tinuteqisaaq sisussat kangaamiut sermiat itillip ilua kangaamiut maniitsoq isortoq alanngua sermersuut kangerlussuatsiaq simiutaq kangerluarsussuaq ikertooq ikertooq sisimiut kuulinnguaq maniitsup sermilia serminnguaq nooralak saqqap kangerluarsua umiiviit maligiaq qaqortorsuaq 52° 67° 67° 66° 54° 66° 54° 50 km 52° fig. 2 the coastal region between isortoq and sisimiut showing place names mentioned in text and key figures. green lines: kangâmiut dykes. nf: nagssugtoqidian front. red stippled line: profile on to which chemical data from dyke and country rock samples from itillip ilua to ikertooq are projected in fig. 97. other symbols in fig. 1. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 5 of 96 geusbulletin.org 2 regional geology the southern extent of the nagssugtoqidian orogen in west greenland was first described based on contrasting structural overprints on the extensive kangâmiut dyke swarm and its host rocks. in a groundbreaking and influential paper, ramberg (1949) noted that towards the north, the progressive structural reworking of the dykes and their archaean host gneisses defined a locally sharp transition into a northern block he called the “nagssugtoqides”, separated from the “kangâmiut gneiss complex” to the south, where the gneisses and the dykes appeared unaffected by nagssugtoqidian overprints. based on extensive reconnaissance mapping, noenygaard & ramberg (1961a) compiled a geological map covering the coastal areas from 63°45n (just south of nuuk) to 69°n (just north of qasigiannguit, see fig. 1). their map shows the broad changes in orientation of the kangâmiut dykes and the structural and metamorphic variations in the host gneisses. south of the nagssugtoqidian, the archaean gneisses are in granulite facies and farther north, these gneisses are, along with the dykes, thoroughly reworked in amphibolite facies. later work (e.g. escher et al. 1975, 1976a, b) further detailed the change in geometry of the kangâmiut dykes (see figs 3, 4) characterising and defining the southern structural limit of the nagssugtoqidian orogen. more recently, the nagssugtoqidian orogen has been subdivided into the southern nagssugtoqidian orogen (sno), the central nagssugtoqidian orogen (cno) and the northern nagssugtoqidian orogen (nno; marker et al. 1995; fig. 1), broadly coincident with respectively the “ikertôq”, “isortôq” and “egedesminde” gneiss complexes of ramberg (1949). the sno is bound by the ikertooq shear zone to the north and to the south by the frontal thrust zone that defines the southern margin of the nagssugtoqidian orogen. the cno extends up to the largescale nordre strømfjord shear zone (bak et al. 1975a, b) and includes tectonically interleaved paleoproterozoic and archaean gneisses and supracrustals in granulite facies. the nno is less penetratively reworked than the sno and cno and displays a gradual transition towards the paleoproterozoic rinkian fold belt in the north (e.g. van gool & marker 2007; thrane 2021). comparable observations in south-east greenland (e.g. bridgwater & gormsen 1968; kalsbeek 1989; nutman et al. 2008; kolb et al. 2016) suggest that the nagssugtoqidian belt extends across greenland and even into the lewisian gneiss complex of nw scotland (e.g. myers 1987; park 1995; mason & brewer 2004). the nagssugtoqidian orogen in west greenland is dominated by archaean gneisses and supracrustals with some paleoproterozoic intrusives and supracrustals, all variably deformed and metamorphosed during the proterozoic. the main crust building episode occurred at 2870–2810 ma and was closely followed by penetrative reworking at 2790–2690 ma (e.g. kalsbeek & nutman 0 200 400 m amphibolite facies gneiss amphibolite dolerite basic agmatite trace of shear plane nf n fig. 3 the classic sketch from escher et al. (1975) showing some of the pertinent features of the nagssugtoqidian front (nf). the dramatic change in dip of dykes, thinning and shearing of the dykes and the reorientation of dykes into parallelism with the dominant gneiss fabric are indicated. escher et al. (1975) based this sketch on field observations in sarfartoq valley (14 km north of kangaamiut sermiat, fig. 1). notably, the ‘front’ is suggested to consist of an array of foreland-directed thrusts. modified from escher et al. (1975). https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 6 of 96 geusbulletin.org 1996; connelly & mengel 2000; connelly et al. 2006). the oldest components identified within the nagssugtoqidian orogen so far are the 3177 ma qorlortoq gneisses (mcintyre et al. 2021; fig. 1). a summary of published geochronological data document the protracted and complex history of magmatism and metamorphism culminating at c. 2850–2750 ma (fig. 5; table 1). minor archaean intrusive activity is recorded in the southern part of the orogen at c. 2500 ma (kalsbeek & nutman 1996; connelly et al. 2000). deformation and metamorphism of these rocks appear to have virtually ceased by c. 2400 ma. accompanying crustal extension and rifting, the mafic kangâmiut dyke swarm (fig. 1) intruded at 2045–2021 ma (kalsbeek & nutman 1996; nutman et al. 1999; connelly et al. 2000; nilsson et al. 2019). the first isotopic dating accomplished on mafic kangâmiut dykes was published by kalsbeek et al. (1978). thirteen samples from two dykes produced a rb–sr isochron age of 1950 ± 60 ma. this date was interpreted to be the emplacement age of those two dykes and was assumed to be representative for all kangâmiut dykes. later, nutman et al. (1999) undertook a u–pb zircon age study of two kangâmiut dykes and detrital zircons from metasediments throughout the nagssugtoqidian region. the first dyke (158078), which was also included in the earlier rb-sr study, produced an age of 2046 ± 8 ma whereas the second dyke (413794) gave an age of 2036 ± 5 ma. this led to the conclusion that the entire kangâmiut dyke swarm was emplaced at c. 2040 ma. at the same time, willigers et al. (1999) published an extensive 40ar–39ar study that included amphibole analyses from 18 dyke samples. they, too, concluded that the primary emplacement age was c. 2040 ma. they fig. 4 distribution and orientation of kangâmiut dykes between kangaamiut and ikertooq. modified from escher et al. 1976b (their fig. 81). the transition across the nagssugtoqidian front (nf) along sarfartoq (middle of figure) is shown in fig. 3. based on field observations and aerial photo interpretation, this figure depicts the dramatic change in orientation of the dykes across the nf. it also highlights at least two distinct orientations of dykes in the foreland region (e–w, nne–ne), mainly east of kangaamiut sermiat, and sporadically in the coastal regions south of itillip ilua. large thrust symbols: thrusts maped by hageskov (1995, 1997, 1998) and olsen (2000) based on fieldwork in the inner sarfartoq valley and helicopter reconnaissance along the nf. their work describes the nf as a series of discrete foreland-directed ductile thrusts rather than one single thrust. small thrust symbols: from the original sketch of escher et al. (1976b), which often coincide with the larger thrust symbols. dashed lines: traces of the nf proposed by noe-nygaard & ramberg 1961a (red), escher et al. 1975 (green) and based on aeromagnetic data (blue; e.g. korstgård et al. 2006; see also fig. 95). ee&w nn&r noe-nygaard & ramberg (1961a) aeromagnetic dataescher et al. (1975)thrust faults a nn&r ee&w a itillip ilua saqqap kangerluarsua kangerlussu aq tasersuaq sarfartoq umiiviit kangaamiut sermiat tasersiaq kangerlussu atsia q km 0 10 20 30 40 50 n f fig. 3 66° 66.5° 53° 52° 51° 66° 66.5° 52° 51° https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 7 of 96 geusbulletin.org also observed younger ages that occur in two clusters with means of 1857 ± 14 ma and 1736 ± 20 ma (fig. 6) and were interpreted to represent “post-peak metamorphic cooling”. a conventional u–pb zircon date was obtained by thermal ionization mass spectrometry (tims) on a dyke by connelly et al. (2000). their result of 2048 +4/–2 ma was considered consistent with previous results. most recently, nilsson et al. (2019) reported a u–pb zircon (tims) analysis from one additional sample (87777) to be 2021 ± 4 ma. these results are further summarised in fig. 6 and table 1. the dykes form a dense, generally nne-trending regional suite of fe-rich tholeiites, extending northwards from isortoq (søndre isortoq) for at least 200 km (fig. 1). the dykes have been found in the sno and the southernmost parts of the cno. similar discordant mafic dykes of possible paleoproterozic age are found in the northernmost part of the cno and the nno (e.g. sørensen 1970, 1971; glassley & sørensen 1980; mengel 1983; àrting 2004; sørensen et al. 2006). however, these dykes have not been dated and their relationship fig. 5 key published age data from the nagssugtoqidian orogen. a: 207pb–206pb age data measured on zircon grains by various methods (sensitive high resolution ion microprobe shrimp; thermal ionization mass spectrometry tims; laser ablation split stream lass, inductively coupled plasma-mass spectroscopy icp-ms). reported 2σ standard deviation (sd) for each analysis is shown. density curve along y-axis illustrates the distribution of the 199 ages. discrete events shown by data clusters are annotated. data from kalsbeek & nutman (1996), whitehouse et al. (1998), nutman et al. (1999), willigers et al. (1999), connelly & mengel (2000), connelly et al. (2000), mcintyre et al. (2021). host rocks kangâmiut dykes calc-alkaline intrusives nagssugtoqidian metamorphism density connelly & mengel 2000 connelly et al. 2000kalsbeek & nutman 1996 mcintyre et al. 2021 whitehouse et al. 1998 1600 1700 1800 1900 2000 2100 2200 2300 2400 2500 2600 2700 2800 2900 3000 3100 3200 3300 a ge (m a) (diameter of dots ~25 ma) a https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 8 of 96 geusbulletin.org with the kangâmiut dyke swarm is currently unknown. in addition to the kangâmiut dykes, the precambrian in west and south-west greenland is cut by several other swarms of proterozoic mafic dykes. some of these are listed together with their referenced ages in table 1. it is notable, that the meta-doleritic, 2053–2029 ma md3 dykes of south-west greenland are broadly the same age as the kangâmiut dykes, despite occurring 200–300 km farther south. an important inference from this is that the extensional tectonic regime responsible for the emplacement of the kangâmiut dykes was clearly not an isolated phenomenon in the western greenland crust c. 2000 myr ago. minor amounts of juvenile, intrusive material was added to the crust between 1950 ma and 1870 ma (‘a’ and ‘s’ in fig. 1; kalsbeek et al. 1987; kalsbeek & nutman 1996; connelly et al. 2000). located along the cno–nno boundary, the field relationships and arc-like calc-alkaline chemical signatures of the arfersiorfik quartz diorite intrusion (‘a’ in fig. 1) was suggested by kalsbeek et al. (1987) to indicate the presence of a cryptic suture along the nordre strømfjord shear zone (fig. 1). more recent work (sørensen et al. 2006; glassley et al. 2014) has further substantiated the presence of such a fundamental tectonic boundary in this part of the nagssugtoqidian orogen. lastly, multi-phase collisional orogeny between 1873 ma and 1775 ma (kalsbeek & nutman 1996; willigers et al. 1999; connelly et al. 2000; van gool et al. 2002) created the nagssugtoqidian orogen. in this study, we fig. 5 (continued) key published age data from the nagssugtoqidian orogen. b: regional map showing the distribution of all dated samples. the samples with metamorphic dates are indicated by the specific metamorphic period recorded. abbreviations in fig. 1. sisimiut charnockite variably reworked archean gneisses supracrustal rocks ma�c dykes kangâmiut ma�c dyke swarm s undi�erentiated archean/paleoproterozoic paleoproterozoic ikertooq sisimiut sisimiut isortuat kangaamiut sermiat n fitilleq kangaamiut maniitsoq nassuttooq kangerl ussu aq nordre isortoq shear zone ikertooq shear zone s c n o sn o iso rto q kangerlussu atsia q greenland ice sheet itillip ilu a 52° 67° 67° 66° 54° 52°54° 1700–1820 ma country rocks 1700–1820 ma kangâmiut dykes 1820–1880 ma country rocks 1820–1880 ma kangâmiut dykes dated sites lacking metamorphic ages b https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 9 of 96 geusbulletin.org focus on the development of the southern part of the orogen and its foreland. the archaean gneisses south of the nagssugtoqidian orogen (referred to as the “kangâmiut gneiss complex” of ramberg 1949) have been characterised by a wealth of field, chronological and isotopic data and the area has been further subdivided into terranes (sensu coney et al. 1980; nelson et al. 2013). for a summary of the developing nomenclature, see e.g. friend & nutman (2019) and steenfelt et al. (2021). detailed aspects of the southern structural front were described in escher et al. (1975). a well-exposed section along sarfartoq valley (fig. 3) shows the dramatic change in dyke orientation characterising the southern nagssugtoqidian front (nf), especially in the areas north-east of the ice cap kangaamiut sermiat (fig. 4). west of the ice cap, the transition is less well defined and the change in orientation of the kangâmiut dykes is more gradual. the continued trace of the southern nf west of kangaamiut sermiat is still a matter of debate. while the structures east of the ice cap are fairly well defined by networks of foreland-directed thrusts, farther west, such structures have not been observed. one reason is likely the inaccessibility of the area between itillip ilua and the area south of the mouth of kangerlussuaq (søndre strømfjord). figure 4 shows three examples of proposed continuation of the nf. on earlier maps (noenygaard & ramberg 1961a), a tentative continuation of the nf was sketched through itillip ilua. however, this was before the itilleq shear zone was known to predate the nagssugtoqidian orogen by more than 600 myr. table 1 geochronological data from selected mafic dykes in west greenland dyke name sample no. location (as named in source) age (ma) method type source maniitsoq to itilleq: kangâmiut dyke 430909 inner søndre strømfjord 1697 ± 15 40ar-39ar (amp) metamorphic willigers et al. 1999 kangâmiut dyke 414915 eastern ikertooq shear zone 1740 ± 15 40ar-39ar (amp) metamorphic willigers et al. 1999 kangâmiut dyke 431205 inner ikertooq 1742 ± 14 40ar-39ar (amp) metamorphic willigers et al. 1999 kangâmiut dyke 430605 inner søndre strømfjord 1745 ± 15 40ar-39ar (amp) metamorphic willigers et al. 1999 kangâmiut dyke 414473 maligiaq 1756 ± 13 40ar-39ar (amp) metamorphic willigers et al. 1999 kangâmiut dyke 430261 n of hamborgerland 1840 ± 80 mineral sm-nda metamorphic stecher et al. 1998 kangâmiut dyke (podded) 431223 n of hamborgerland 1863 ± 13 40ar-39ar (amp) metamorphic willigers et al. 1999 kangâmiut dyke (podded) 431223 n of hamborgerland 1873 ± 13 40ar-39ar (amp) metamorphic willigers et al. 1999 kangâmiut dyke 158078 nnw of kangaamiut 1950 ± 60 rb-sr (wr) igneous kalsbeek et al. 1978 kangâmiut dyke (ene) 448916 simiutak, mouth of søndre strømfjord 1981 ± 18 40ar-39ar (amp) metamorphic willigers et al. 1999 kangâmiut dyke (e-w) 448933 kangerluarsugssuaq 1981 ± 26 40ar-39ar (amp) metamorphic willigers et al. 1999 kangâmiut dyke (ene) 431264 mouth of søndre strømfjord 2013 ± 14 40ar-39ar (amp) igneous? willigers et al. 1999 kangâmiut dyke (ene)) 448915 kangerluarsugssuaq 2021 ± 13 40ar-39ar (amp) igneous? willigers et al. 1999 kangâmiut dyke 87777 n of hamborgerland 2021 ± 4 u-pb (zrn) igneous nilsson et al. 2019 kangâmiut dyke (undeformed) 431229 inner itilleq 2028 ± 29 40ar-39ar (amp) igneous? willigers et al. 1999 kangâmiut dyke 413794 outer itilleq 2036 ± 5 u-pb (zrn) igneous nutman et al. 1999 kangâmiut dyke 158078 nnw of kangaamiut 2046 ± 8 u-pb (zrn) igneous nutman et al. 1999 kangâmiut dyke 430248 n of hamborgerland 2048 +4/–2 u-pb (zrn) igneous connelly et al. 2000 kangâmiut dyke 87777 n of hamborgerland 2478 ± 7b 40ar-39ar (amp) igneous? nilsson et al. 2019 kangâmiut dyke 87777 n of hamborgerland 2492 ± 7b 40ar-39ar (amp) igneous? nilsson et al. 2019 kangâmiut dyke 87777 n of hamborgerland 2528 ± 25b 40ar-39ar (amp) igneous? willigers et al. 1999 paamiut to nuuk: md2 (ne) 511577 sermilik, e of paamiut 2209 ± 5 u-pb (bdy) igneous nilsson et al. 2013 md3 468749 sermilik 2029 ± 3 u-pb (bdy) igneous nilsson et al. 2010 md3 515830 inner bjørnesund 2040 ± 3.1 u-pb (bdy) igneous nilsson et al. 2010 md3 515805 ravn storø 2050 ± 2 u-pb (bdy) igneous nilsson et al. 2010 md3 (e-w) 508242 inner bjørnesund 2053 ± 2 u-pb (bdy) igneous nilsson et al. 2013 md3 (n-s) 508210 inner bjørnesund 2049 ± 6 u-pb (bdy) igneous nilsson et al. 2013 md3 (wnw) 508213 inner bjørnesund 2042 ± 2 u-pb (bdy) igneous nilsson et al. 2013 bn (n-s) isukasia 2214 ± 10 u-pb (zrn) igneous nutman et al. 1995 e-w dykes 508214 inner bjørnesund 2365 ± 2 u-pb (bdy) igneous nilsson et al. 2013 e-w dykes 508201 grædefjord 2374 ± 4 u-pb (bdy) igneous nilsson et al. 2013 aplagioclase, amphibole and garnet (metamorphic minerals). bindicating excess ar, hence geologically meaningless (nilsson et al. 2019). amp: amphibole. bdy: baddeleyite. wr: whole rock. zrn: zircon. location names as per source: søndre strømfjord (kangerlussuaq), hamborgerland (sermersuut), kangerluarsugssuaq (kangerluarsussuaq, sisimiut), itilleq (itillip ilua), bjørnesund (allumersat), ravn storø (takisup qeqertarsua), grædefjord (kangerluarsussuaq (nuuk)). https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 10 of 96 geusbulletin.org escher et al. (1975) proposed an arcuate boundary that also ended up near itillip ilua (fig. 4). the aeromagnetic anomalies show a boundary that appears to continue towards the south-west (fig. 4). it is possible that more distributed shortening across the nf was accommodated in the ductile regime in the coastal regions where deeper crustal levels are exposed, and that shortening at the shallower crustal levels exposed towards the greenland ice sheet manifested itself as recognisable thrusts. regardless of the favored interpretation, any model is, at present, highly speculative due to the lack of structural observations in the critical, but highly inaccessible areas west of kangaamiut sermiat. while the emplacement of the kangâmiut dykes is now known to pre-date the nagssugtoqidian deformation and metamorphism by at least 150 myr, the relationship between dyke intrusion and pre-existing structures is still a topic of some debate. in the inussuttusup tunua area (anders olsen sund), escher et al. (1975) and jack (1978) argued that the dykes were emplaced into an overall contractional environment. the ne-trending kangâmiut dykes occurred typically in sinistral shear zones, whereas e–wto ese–wnw-trending dykes were located in dextral shear zones striking c. 100°. this conjugate set of shear zones was interpreted to represent a nnw–sse-orientated maximum horizontal stress, and the authors further argued, based on field interpretations, that dilational opening took place parallel to the minimum horizontal stress direction (ene–wsw). the conjugate set of fractures was believed to partly pre-date dyke emplacement and to have been formed partly or reactivated by elevated hydrostatic magma pressures during dyke emplacement. nash (1979a, b) suggested that similar field relationships could be seen between pre-existing structures, dyke emplacement and orientation of the accompanying palaeo-stress system in the western part of itillip ilua. recent dating of kangâmiut dykes have, however, necessitated alternative models for the observed structural relationships (see section 3.1). fig. 6 all published dates for kangâmiut dykes, sorted by age. diameter of dots c. 10 ma. also shown are the computed means and 1 sd for the apparent clusters in the ranges c. 1760–1700 ma (blue shading), c. 1870–1840 ma (orange shading) and c. 2070–1940 ma (green shading). density curve is shown along the y-axis. the reported 2 sd for each age is shown as vertical bars. data from kalsbeek et al. (1978), kalsbeek & nutman (1996), stecher et al. (1998), nutman et al. (1999), willigers et al. (1999), connelly et al. (2000) and nilsson et al. (2019). ar–ar amp: 40ar–39ar plateau ages from amphiboles. rb–sr wr: whole-rock isochron age from two kangâmiut dykes. shrimp zrn and tims zrn: 207pb–206pb ages from zircons by shrimp or tims. sm–nd mins: mineral isochron age based on amphibole, garnet and plagioclase from kangâmiut dykes. diameter of dots c. 10 ma 1600 1700 1800 1900 2000 2100 2200 a ge (m a) ar-ar amp rb-sr wr shrimp zrn sm-nd mins tims zrn mean = 2035 ma sd = 17 ma mean = 1857 ma sd = 14 ma mean = 1736 ma sd = 20 ma density https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 11 of 96 geusbulletin.org based on field observations in the area east of kangaamiut sermiat, hageskov (1995, 1997, 1998) and olsen (2000) argued that the kangâmiut dyke swarm did not intrude along pre-existing faults or shear-zones, in contrast to the syn-intrusive conjugate shear system proposed above. rather, they argued that during emplacement, the dykes generated their own extensional fractures opening perpendicular to both local and regional minimum horizontal palaeo-stress fields. the age and nature of the metamorphic overprint displayed by the kangâmiut dykes are also a topic of discussion. several earlier descriptions state that in the orogenic foreland to the south, dykes are generally unmetamorphosed and undeformed (i.e. ‘fresh’) and that it is only north of itillip ilua that dykes display any significant metamorphism (e.g. ramberg 1949; noe-nygaard & ramberg 1961b). in apparent contrast, works by windley (1970) at maniitsoq and jack (1978) at inussuttusup tunua, just south-west of itillip ilua, describe metamorphic overprints on the igneous mineral assemblages of the dykes. however, these studies attributed the overprints to late-magmatic features associated with so-called ‘autometamorphic’ mineral growth during elevated temperatures, immediately following emplacement. subsequently, korstgård (1980) interpreted all metamorphic overprints between itillip ilua and ikertooq as being of nagssugtoqidian age. more recent reconnaissance work between itillip ilua and maniitsoq by the danish lithosphere centre (e.g. mengel et al. 1997, 1998) mapped the extent of such metamorphic overprints to >100 km south of itillip ilua, and – albeit based on a limited data set – also argued that this metamorphic overprint was of nagssugtoqidian age. dating of metamorphic minerals (willigers et al. 1999; stecher et al. 1998) supported this suggestion. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 12 of 96 geusbulletin.org 3 dykes and country rocks from maniitsoq to maligiaq in this chapter we outline the field relations of the dykes and their host rocks in the study area. from maniitsoq to itillip ilua (figs 1, 2; section 3.1), the dykes and their host gneisses suffered negligible structural overprints due to the nagssugtoqidian orogeny, whereas we will argue that the dykes record a range of metamorphic effects. in contrast, between itillip ilua and maligiaq (figs 1, 2; section 3.2), the nagssugtoqidian structural and metamorphic overprints are pronounced and penetrative in both dykes and their country rocks. several generations of dykes occur in the area studied here. by far the volumetrically dominant is a suite of nne–ne-trending mafic dykes, which are traditionally considered to belong to the main kangâmiut dyke swarm, and which are the main focus of the present study. e–w-trending dykes occur throughout but are much less common. mayborn (2000) identifies two compositionally different suites of e–w dykes: orthopyroxene-bearing “ob dykes” and plagioclase-phyric “pp dykes” and finds that the main kangâmiut dyke swarm generally crosscuts both types. similarly, windley (1970) and jack (1978) describe e–w dykes from the maniitsoq and inussuttusup tunua areas, respectively (sections 3.1.2 and 3.1.5) predating the main kangâmiut dyke swarm. it is important to emphasise that there are currently no age constraints on these potentially different e–w-trending dykes. consequently, we will use the term ‘kangâmiut dykes’ to describe them all while acknowledging that further studies may well suggest the existence of temporally different intrusive events. 3.1 dykes and country rocks from maniitsoq to itillip ilua the regional extent of the kangâmiut dyke swarm is well-known from early reconnaissance work (e.g. noenygaard & ramberg 1961a) and members of the swarm can be found intruding archaean gneisses as far south as c. 10 km south of the mouth of isortoq (fig. 2). 3.1.1 isortoq to maniitsoq around isortoq, host rocks are polyphase tonalitic to granodioritic gneisses in amphibolite facies (fig 7a, b), locally hosting concordant supracrustal and mafic units. planar fabrics in the gneisses are generally subvertical and ne-striking. kangâmiut dykes are typically less than a few metres wide with sharp contacts along which occur intricate small-scale intrusive features, bridges, and chilled contact zones (fig 8a–c). the dykes generally show right-stepping asymmetry (fig. 8a). coarsegrained leucocratic pegmatites, with conspicuous black feldspars (fig. 7b), truncate the gneissic host rocks and are themselves cut by kangâmiut dykes. alanngua (fig. 2) appears to mark the southernmost occurrence of kangâmiut dykes. the c. 3 ga finnefjeld orthogneiss complex (gardiner et al. 2019) occurs immediately east of alanngua and was not found to be intruded by kangâmiut dykes. fig. 7 field examples of dykes. a: archaean grey tonalitic to granodioritic polyphase layered gneisses in amphibolite facies, alanngua (view towards sse). notebook (in yellow dashed circle) is 11 x 18 cm. b: same as panel a but showing detail of discordant, coarse-grained, undeformed plagioclase-rich pegmatite. c: migmatised metapelitic unit in granulite facies gneisses from north-east sermersuut. large reddish garnets and smaller brownish orthopyroxenes are visible in coarser-grained layers. scale is 10 cm. a b c https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 13 of 96 geusbulletin.org from isortoq and towards maniitsoq (fig. 2), the host gneisses are in granulite facies, as shown by orthopyroxene in leucocratic patches in both gneisses and in schistose mafic and felsic supracrustals (fig. 7c). overall, the planar fabrics in the gneisses around maniitsoq are subvertical with variable strikes reflecting large-scale folding. in this area, the kangâmiut dykes trend nne and, like further south, preserve intrusive features and show right-stepping asymmetry. locally, a weak margin-parallel foliation is developed in the dykes. on the eastern side of sermersuut (hamborgerland; fig. 2), dykes cut across migmatised supracrustals and display subtle sinistral drag within their margins. also notable is that dykes here show left-stepping intrusion geometries as opposed to the dominantly right-stepping dykes farther south (fig. 8b). along the north shore of isortoq, subtle metamorphic overprints on the igneous dyke mineralogy are visible in outcrop, as igneous clinopyroxenes are partially replaced by amphiboles and garnets are often observed. the proposed extension of the aujassoq–evighedsfjord shear zone (allaart & jensen 1979) projects through the kuulinnquaq islands, just a few km se of maniitsoq (fig. 2; cf. 1:500 000 geological map, allaart 1982). on one of these, an island c. 100 m across, the shear belt is characterised by alternating panels of protomylonite, schistose zones and normally foliated gneiss. their axial planar foliations are subvertical and strike ne, and lineations – defined by amphibole and aggregates in mafic pods – plunge shallowly to moderately sw. all observed kinematic indicators (delta-clasts) on subhorizontal surfaces indicate sinistral shear. farther north-east of kuulinnquaq, across smaller islands and along the shores of maniitsup sermilia (fig. 2), steeply dipping gneissic structures are consistently ne-trending and define a distinct zone of parallel planar fabrics. allaart & jensen (1979) noted that farther inland, dykes appear unaffected by the linear belt, thus suggesting a pre-dyke age for this major structure. older dykes are occasionally also observed. one mafic, discordant and overall e–w-trending dyke with 1–2 cm black feldspar phenocrysts is folded, contains the same planar ne-striking fabric as the country gneisses and is assumed to be of pre-kangâmiut age. fig. 8 field examples of dykes. a: 10 cm wide, discordant, undeformed kangâmiut dyke, hosted by archaean grey tonalitic gneisses in amphibolite facies. dyke shows delicate intrusive features (hooks, bridges, bayonets) consistent with right-stepping offsets, alanngua. scale is 10 cm. b: detail of the margin of a 50 cm wide kangâmiut dyke that cuts granulite facies host gneisses discordantly. a left-stepping offset is indicated by a broken bridge. subtle drag folding of possible cooling joints within the dyke’s horn is consistent with minor sinistral displacement along this margin. north-east sermersuut. scale is 10 cm. c: a 4 m wide kangâmiut dyke that discordantly cuts the layering of grey, tonalitic gneisses in amphibolite facies. narrow abandoned fracture in left foreground shows overall left-stepping intrusion geometry. sulussugut (finnefjeld) in background. alanngua. notebook at arrow is 11 x 18 cm. a c b https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 14 of 96 geusbulletin.org just ne of maniitsoq, an ese-trending, 2 m wide, brown, discordant and vesicular dyke, with a pronounced sidewall-parallel layering, was noted. its relationship with the generally nne-trending kangâmiut dykes was, however, not exposed. 3.1.2 maniitsoq to kangerlussuatsiaq between maniitsoq and the kangerlussuatsiaq (evighedsfjord), dykes maintain their nne trend but are wider and the swarm appears overall denser. figure 9 shows a typical occurrence of nne-trending dykes across parts of an island south of the mouth of kangerlussuatsiaq. host gneisses and associated supracrustal units are all in granulite facies and exhibit folding at various scales – unlike those farther south. there are also some structurally simpler granitic intrusions, which are cut by kangâmiut dykes and assumed to be of late archaean age. in the area around maniitsoq, windley (1970) finds that the nne-trending kangâmiut dykes truncate two older dyke sets as (1) a ne-trending set of black ophitic dolerites, and (2) a se-trending set of ‘blue’ porphyritic dolerites. in addition, the kangâmiut dykes are themselves cut by se-trending, brown-red and vesicular dykes. the exact ages of these other dykes are not known. the wider kangâmiut dykes are often compositionally zoned, from fine-grained and ophitic to sub-ophitic chilled margins, to gradually coarser grained and plagioclase-rich central parts (fig. 10a). garnets are common in the more felsic centres along with amphibole and plagioclase. in a detailed study of dykes between north and north-east of sermersuut (fig. 2), it was proposed that the largely symmetrical zoning of the larger dykes was caused by a single protracted intrusive event (windley 1970). this event led to earlier magma injections that chilled rapidly and preserved an igneous mineralogy, whereas subsequent pulses involved a more felsic (and hydrous) magma, which led to the metamorphic garnet-amphibolites. the dykes between maniitsoq and kangerlussuatsiaq exhibit a range of metamorphic overprints observable in outcrop, including amphiboles replacing clinopyroxenes and the rare to widespread presence of garnets. 3.1.3 kangerlussuatsiaq to simiutaq at the mouth of kangerlussuatsiaq (fig. 2), there are several examples of cross-cutting dykes (figs 10c, 11a). on one of the smaller islands, an e–w dyke is cut by a ne-trending dyke with a chilled margin against the earlier dyke reflecting the relative age relationships seen elsewhere (fig. 11b). between kangaamiut and simiutaq, at the mouth of kangerlussuaq, overall dyke trends begin to subtly bend from nne towards ne (fig. 2) as the swarm also appears slightly denser. host gneisses are all in granulite facies, with planar fabrics with varying strikes and moderate to steep northerly to westerly dips. ungusivik island (midway between kangaamiut and simiutaq; fig. 2) is where one of the first dated kangâmiut dykes was collected (age presented by kalsbeek et al. 1978). the 1950 ± 60 ma rb–sr age (see also chapter 9) was obtained from samples on ungusivik and at sisussat (fiskemesterens havn), on the central, leucocratic portions of nne-trending, c. 60 m wide composite dykes, cutting granulite facies gneisses. later, nutman et al. (1999) obtained a 2046 ± 8 ma u–pb zircon shrimp age from the dyke on fig. 9 a dense swarm of nne-trending kangâmiut dykes (yellow dashed lines) on the island itilliup qaqqaa at the mouth of kangerlussuatsiaq. key dykes have been outlined for clarity. the dyke slightly right of the centre of the image (arrow) is c. 40 m wide. layering in host gneisses strike ene. image from google earth (centre of image at 65°44’10n, 53°14’30w). location in fig. 2. 500 m background image from google earth fig. 11bn https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 15 of 96 geusbulletin.org ungusivik. the dyke has a well-developed, steep, margin-parallel foliation, with both swand ne-plunging stretching lineation defined by amphibole aggregates. kinematic indicators on horizontal surfaces show consistent sinistral shear. metamorphic garnets are developed in most dykes, particularly in the leucocratic centres of wider dykes. the same range of metamorphic overprints can also be found throughout the coastal dyke exposures between kangaamiut and simiutaq. at sisussat, just south of simiutaq (fig. 2), the second of the two dykes dated by kalsbeek et al. (1978) was sampled. at this locality, the dykes are up to 60 m wide, typically composite with leucocratic, garnet-bearing centres. they are overprinted by steeply dipping and nne-striking foliated zones containing shallow ssw-plunging stretching lineation, defined by amphibole aggregates. on horizontal surfaces, kinematic indicators always show sinistral asymmetry. on simiutaq island, kangâmiut dykes are ubiquitous and trend systematically 25–30°. the dykes are typically up to 40–60 m wide and occur about every 500–1000 m. figure 12 shows the north coast of simiutaq and the regular spacing of kangâmiut dykes as well as an earlier formed 70 m wide e–w-trending dyke (fig. 13a). host rocks are granulite facies gneisses with abundant mafic inclusions and layers. the dykes have preserved a range of delicate, brittle intrusive features, such as bayonets (fig. 13b). where these are developed, dykes are left-stepping, similar to what was observed along the coast from the north shore of sermersuut and farther north. the metamorphic overprints observed here are the same as farther south: amphiboles replacing clinopyroxenes and rare to widespread garnets in leucocratic parts of the dyke. entering the nagssugtoqidian frontal thrust zone from kangerlussuaq, the curvature of the dyke swarm becomes increasingly apparent. only 30–40 km east of simiutaq, the dykes mainly strike ne to ene parallel to the kangerlussuaq fjord, and dip gently nw to nnw (figs 4, 14). at serminnguaq (fig. 4), hansen (1989) notes that the host rocks are amphibolite facies granodioritic gneisses and that metamorphic overprints on the dykes can be seen in outcrop. further into kangerlussuaq, opposite sarfartoq (fig. 2), mayborn (2000) found that dykes are generally foliated but preserve intrusive relationships with amphibolite facies gneisses that are interpreted as reworked and retrogressed archaean granulite facies gneisses (fig. 14). 3.1.4 kangerluarsussuaq further north along the outer coast from simiutaq, the kangerluarsussuaq fjord (fig. 2) offers an excellent 25 km e–w-orientated section through the dyke swarm. mayborn (2000) investigated this section in detail, and the following description is based on that work. dykes trend on average 30° and vary in width from less than 1 m to 140 m. excluding the widest (140 m) and thinnest (0.4 m), the mean width is c. 30 m. the host gneisses are in granulite facies throughout. the compositional layering in the gneisses generally strike in a north-westerly direction with moderate to steep south-westerly dips. thinner dykes and the margins of wider dykes only show incipient amphiboleand garnet-growth. in contrast, wider dykes display a zoning similar to that described above and by a b c strain gradient-> nne-d yke ne-d yke fig. 10 field examples of dykes. a: coarse-grained garnet-bearing amphibolite from centre of c. 40 m wide zoned dyke. akuliaruseq, ne of sermersuut (location in fig. 2). scale is 10 cm. b: margin of 1.5 m wide nne-trending kangâmiut dyke, showing localised foliation development in 0.2–0.5 m wide margin-parallel zones. north shore of tunu (hamborgersund) north of sermersuut. scale is 10 cm. c: a 30 m wide nne-trending kangâmiut dyke cut by 0.5 m wide ne-trending dyke. thin chilled margin in ne-trending dyke (arrow). tinuteqisaaq, a small island on south side of the mouth of kangerlussuatsiaq. scale is 10 cm. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 16 of 96 geusbulletin.org windley (1970), with leucocratic centres showing strong foliation and a completely recrystallised mineral assemblage of plagioclase, amphibole, garnet and biotite. inside a 140 m wide dyke, the central leucocratic component appears to intrude the marginal, less altered parts of the dyke as late-stage metamorphic epidote-quartz-feldspar veins (mayborn 2000). similar features were observed in the wide dyke south of simiutaq (section 3.1.3). a ne-trending, orthopyroxene-bearing dyke shows the same metamorphic overprints as ‘typical’ kangâmiut dykes, while an e–w-trending, plagioclase-phyric 2 m wide dyke was observed near the mouth of kangerluarsussuaq and showed only minor metamorphic overprints. 3.1.5 inussuttusup tunua area from the stretch of coast between kangerluarsussuaq and itillip ilua, most observations come from the work of jack (1978) in the inussuttusup tunua area (fig. 2). the country rocks are granulite facies tonalitic to granodioritic gneisses, with only minor supracrustal material, notably in the ne inussuttusup tunua area. here, mixed supracrustal units and gneisses occur in the core of a 10 km long by 3 km wide ne-trending structural basin, surrounded by reworked or retrogressed gneisses in amphibolite facies (davidson 1978; jack 1978). the north-western half of the inussuttusup tunua area is characterised by flat-lying structures, only interrupted in the south-west by a c. 2 × 4 km dome structure, rimmed by supracrustals and cored by granite (davidson 1978; jack 1978). the dykes truncate all of the above units with dominant trends of 30–35° in the southern part of the inussuttusup tunua area and up to 45° in the north, towards the mouth of itillip ilua (fig. 15). in a detailed study of the dykes in the area, jack (1978) described several sets. type 1 are smaller (<10 m) nne-trending dykes. type 2 are larger (>10 m) and a b 0.5 m fig. 11 field examples of dykes on itilliup qaqqaa, a small island on the south side of the mouth of kangerlussuatsiaq. a: contact (yellow dashed line) between a 70 m wide, nne-trending coarser-grained kangâmiut dyke (beneath) and a later, finer-grained, ne-trending dyke (above). the later dyke has a subtle, chilled margin against the coarse-grained dyke and displays consistent left-stepping intrusion geometry on horizontal surfaces. scale is 10 cm. b: 30 m wide nne-trending kangâmiut dyke, cut by ne-trending c. 0.5 m wide dyke located in the north-western corner of itilliup qaqqaa (location in fig. 9). 500 m fig. 13a background image from google earth n fig. 12 satellite view of a 70 m wide e–w dyke, cut by a swarm of nne-trending kangâmiut dykes. at this location, nne dykes occur approximately every 400–500 m and are up to 40–50 m wide. some of the larger dykes are outlined (yellow dashed lines). north shore of simiutaq island, in the mouth of kangerlussuaq. image from google earth. small island in centre of image is at 66°05’16n, 53°33’10w. location in fig. 2. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 17 of 96 geusbulletin.org substantially more abundant ne-trending dykes. type 3 are e–w-trending doleritic dykes and type 4 are e–w-trending black feldspar porphyritic dykes. type 4 dykes appear to be the earliest, whereas cross-cutting relationships between types 1, 2 and 3 suggest that these are largely contemporaneous. the overall density of dykes increases from northwest to south-east in the area studied. jack (1978) further found that basement rocks are cut by two sets of conjugate shear zones, with a 100–110°-trending set of narrow and discrete dextral cataclastic zones and an ne-trending set of wider and sinistral ductile shear zones. most dykes typically show apophyses and bayonets that display a systematic sense of asymmetry for individual swarms (figs 2.5, 2.6 in jack 1978), where e–w-trending dykes show left-stepping offsets and ne-trending dykes show right-stepping offsets. jack (1978) argued that the host rocks’ conjugate shear zones also exerted some control on dyke geometry and that the highest densities of dykes coincide with the best developed ne fabrics in the gneisses. figure 16 illustrates the pronounced correlation between dyke orientation and the strong gneiss fabric developed in the gneisses. it is clear that the inussuttusup tunua area represents a key transition between overall nne-trending dykes along the coast farther south, and the itillip ilua area where dykes are largely parallel to the dominant e–w-structures in the host rocks. in terms of metamorphic overprints, the dykes around inussuttusup tunua show the range from marginal samples displaying only minor overprints to samples from dyke centres, where the rock is totally recrystallised to a foliated garnet-amphibolite. with few exceptions, the original igneous texture is nevertheless still preserved or recognisable. in some dykes, garnet growth appears to have been initiated along irregular veins, then migrated into the rock and resulted in dendritic or ‘cauliflower’ textures. primary orthopyroxene is found in a few of the dykes, whereas primary amphibole has not been positively identified but was posited by hansen (1989) from the nooralak peninsula just ne of inussuttusup tunua in outer itillip ilua (fig. 2). jack (1978) described a sequence of metamorphic reactions, involving amphibole or biotite rims on ilmenite or clinopyroxene, or both, and a later stage of garnet growth that partially or completely replaced these rims. these observations correspond closely to those outlined in this study in chapter 5. jack (1978) attributes these corona-forming reactions to the mechanism of ‘autometamorphism’, which purports that metamorphic reactions progressed immediately after dyke emplacement driven by elevated temperatures and increased fluid content associated with later magma pulses. we propose an alternative interpretation of the initiation, progress and cause of these metamorphic reactions (see section 3.1.6). 3.1.6 itillip ilua and surrounding areas in inner itillip ilua and along the north shore of the outer fjord, dykes are broadly parallel with a predominant e–w-trending layering and foliation of the gneisses. away from the fjord, dykes trend ene to ne (fig. 16). at first glance, it appears to be a simple case of dykes and their host rocks being overprinted by a dextral shear a b <e–w dyke -> <n n e–ssw dyke -> fig. 13 field examples of dykes on simiutaq island, at the mouth of kangerlussuaq. a: 70 m wide e–w dyke cut by 0.5 m nne-trending kangâmiut dyke. location in fig. 12. b: discordant nne–ne-trending kangâmiut dyke showing left-stepping intrusion features. host rocks are heterogeneous, migmatised amphibolite facies gneisses. north shore of simiutaq. hammer (in yellow dashed circle) is 40 cm long. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 18 of 96 geusbulletin.org zone in which all structures and intrusions were rotated into parallelism by the stress-field responsible for the shear zone. based on field observations, the structural history along and near itillip ilua is substantially more complicated. while it may be debatable whether the strong planar fabrics along itillip ilua conform to fig. 14 view ene from the north shore of kangerlussuaq towards sarfartoq (north wall of sarfartoq valley in the background). centre of image shows a pair of 1 m wide kangâmiut dykes (indicated by arrows) thoroughly rotated into parallelism with nnw-dipping layering in deformed amphibolite facies host gneisses. locally, the discordant nature of the dykes is preserved (e.g. at tip of the uppermost arrow). 200 m background image from google earth n fig. 15 satellite view of ne-trending dykes cutting an ese-trending dyke on the north coast of inussuttusup tunua. individual dykes in both sets are c. 35–40 m wide in this image. key dykes are outlined (yellow dashed lines) for clarity. image from google earth. intersection of dykes located at 66°31’34n, 53°27’10w. location in fig. 2. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 19 of 96 geusbulletin.org common definitions of a shear zone (e.g. ramsay 1980), we nonetheless use the term ‘itilleq shear zone’ here (fig. 16). earlier workers observed that kangâmiut dykes cut gneisses with a strong shear fabric in the itillip ilua area and described the dykes as having only minor structural and metamorphic overprints. farther north, in ikertooq, however, the dykes and their country rocks are strongly deformed and metamorphosed in amphibolite and granulite facies. in order to distinguish between these events, watterson (1974) used the terms “nagssugtoqidian 1” (nag. 1) and “nagssugtoqidian 2” (nag. 2). nag. 1 described the tectonic event responsible for the strong planar fabrics in the itilleq shear zone and similar structures cut by dykes. nag. 2 described a later high grade deformational and metamorphic event, which created the ikertooq shear zone and similar structures throughout the nagssugtoqidian orogen. subsequent rb–sr whole-rock geochronology, however, indicated that nag. 1 most likely was an archaean event (and thus unrelated to nagssugtoqidian deformation and metamorphism), which led kalsbeek & zeck (1978) and kalsbeek (1979) to recommend abandoning the nag. 1 and nag. 2 terminology. the strong planar fabric displayed in the itilleq shear zone is truncated by tonalitic, dioritic, and granitic bodies in several places along itillip ilua (figs 17a–c). a few of these undeformed rocks were dated by u–pb of zircons to around 2.5 ga (2492 ± 12 ma by kalsbeek & nutman 1996, and 2498 ± 4 ma by connelly & mengel 2000), confirming that the itilleq shear zone is an archaean feature. an identical rb–sr age (2500 ma ± 40) was obtained by hickman (1979) on an “aplite associated with a latenag. 1 pegmatite near itilleq”. it is important to remember, that c. 450 my separate the older straight belt fabrics in the itilleq shear zone from the events leading to the intrusion of kangâmiut dykes and accompanying deformation. also, 150 my n w s e 30 15021 0 240 300 330 n w s e 30 60 120 15021 0 330 n s e 30 60 120 150 w 21 0 240 300 330 n w s e 30 60 120 15021 0 240 300 330 n w s e 30 60 120 15021 0 240 300 330 n w s e 30 60 120 15021 0 240 300 330 n w s e 30 60 120 15021 0 240 300 330 n w s e 30 60 120 15021 0 240 300 330 n w s e 30 60 120 15021 0 240 300 330 n w s e 30 60 21 0 300 330 n w s e 30 60 120 15021 0 240 n w s e 30 60 120 15021 0 240 300 330 (n = 48)(n = 97) (n = 147)(n = 582) (n = 695) (n = 146) (n = 13)(n = 57) (n = 40) (n = 38) n = 40 (n = 41) west east gneiss and dyke trends, qaqqatoqaq dyke trends, west itillip ilua foliation/shear zones in it dyke trends, it planar measurements gneisses, south shore itillip ilua dyke trends, east itillip ilua gneisses, north shore itillip ilua linear measurements dyke orientations qaqqatoqaq nuussuup qulaaitillip ilua itillip ilua itilleq inussuttusup tunua 66°30´53°00´ n 5 km eqalugaarsuit nuussuaq nooralak itilleq shear zone granulite facies gneiss amphibolite facies gneiss anorthosite supracrustal gneiss amphibolite dyke fig. 16 orientation plots for dykes and host gneisses in areas along and adjacent to itillip ilua. data from the inussuttusup tunua (it) area are from jack (1978), the remainder are from this study. see fig. 38 for contouring details. planar measurements (blue shading) include all penetrative planar elements on outcrop scale in metadykes and host gneisses, typically a combination of mineral foliation and compositional layering (e.g. fig. 22). linear measurements (red shading) include all penetrative linear elements in metadykes and host gneisses such as mineral lineation, stretching lineation (typically quartz and feldspar aggregates), rodding, mullions and axes of minor folds (e.g. fig. 23). green dashed lines: denote the itilleq shear zone. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 20 of 96 geusbulletin.org separate that event from the earliest nagssugtoqidian orogenic activities (see also chapter 9). from inussuttusup tunua and north, it appears that host rock structures and dyke orientations gradually bend clockwise, parallel to the predominant trend around itillip ilua (fig. 16). in the following sections we describe the rocks and structures along itillip ilua, outside (section 3.1.6.1) and inside (section 3.1.6.2) the itilleq shear zone. 3.1.6.1 qaqqatoqaq and nuussuup qulaa the country rocks at qaqqatoqaq and nuussuup qulaa (fig. 16) are believed to have suffered their last major phases of deformation and metamorphism in the archaean (2740–2710 ma, kalsbeek & nutman 1996). the gneisses on qaqqatoqaq and nuussuup qulaa are intruded by the kangâmiut dykes, which crosscut all structures and lithologies of the archaean rocks. the bulk of the country rocks are quartzo-feldspathic a b c fig. 17 field examples of discordant tonalitic bodies from locations in and around itillip ilua. a: vein of grey tonalite cutting across strong planar fabric in amphibolite facies host rocks. south shore, inner itillip ilua. scale is 10 cm. b: irregular, discordant body of grey tonalite with angular fragments of a strongly deformed and banded mafic and felsic host amphibolite facies gneiss. brittle injections of delicate tonalite veins and off-shoots riddle the host gneiss in the lower left part of the panel (just left of ruler scale). south shore, inner itillip ilua. scale is 10 cm. c: grey tonalite vein, discordantly truncating the strong planar fabric of an amphibolite facies host gneiss. apparent dextral offset along margin of intrusive vein can be seen in agmatite zone in centre of image. delicate intrusion features show right-stepping geometry in a largely brittle domain. north shore, head of itillip ilua. scale is 10 cm. a b fig. 18 field examples of dykes from saqqap kangerluarsua. a: veined grey gneiss on approaching a shear zone. hammer is 45 cm. b: grey gneiss veined by leucocratic material. hammer in yellow dashed circle is 45 cm. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 21 of 96 geusbulletin.org gneisses of mainly tonalitic to granodioritic composition. the gneisses are heterogeneously layered rocks often transected by leucocratic veins forming an irregular network (fig. 18a). mafic inclusions of various shapes are distributed throughout the gneisses and occur as isolated lenses, pods and agmatite zones (fig. 18b). rocks of obvious supracrustal affinity are rare at qaqqatoqaq and generally do not constitute mappable units. only one isolated occurrence of garnet-sillimanite-biotite gneiss was found in the central part of qaqqatoqaq associated with a layered amphibolite. at the head of itillip ilua, a large mass of anorthosite is present (ellitsgaard-rasmussen & mouritzen 1954). the anorthosite is deformed and intruded by kangâmiut dykes (fig. 19a–c). in fig. 19d, the anorthosite is cut by a gray tonalitic-dioritic sheet, which is remarkably similar to the c. 2.5 ga tonalitic-dioritic bodies mentioned previously (section 3.1.6; fig. 17). while the tonalite-diorite sheet at this location has not been dated, its similar composition and appearance would support a pre-2.5 ga age for the anorthosite. the quartzo-feldspathic gneisses and their mafic inclusions show granulite facies assemblages in the western parts of qaqqatoqaq and nuussuup qulaa, and amphibolite facies assemblages in the eastern parts (fig. 16). no large-scale structures have been mapped on qaqqatoqaq and nuussuup qulaa, mainly due to an absence of appropriate marker horizons. further west, in the inussuttusup tunua area, a major basin and a smaller dome structure, outlined by supracrustal rocks, were mapped by davidson (1978) and jack (1978; section 3.1.5; fig. 16). the planar fabric in the quartzo-feldspathic gneisses is a combination of coarse compositional layering and a preferred habit orientation of mainly biotite (schistosity). on qaqqatoqaq, the general structural trend, shown by the planar fabric, is moderately to steeply neor sw-dipping (fig. 16), whereas on nuussuup qulaa, south of itillip ilua, the dominant strike is ne–sw. a gradual clockwise rotation of structures on approaching itillip ilua from the south is also evident (fig. 16). in both areas, the challenging topography means that ground observations are limited. ba c d fig. 19 field examples of dykes and anorthosite from the head of itillip ilua. a: view towards east of the far western extent of the anorthosite body (qaqortorsuaq mountain; ellitsgaard-rasmussen & moritzen 1954). amphibolite facies host gneisses in left foreground. in the centre of the photo, two 10–20 m wide, e–w-trending and discordant kangâmiut dykes (two yellow arrows), can be traced up the hill, turning slightly to the right (south). another, wider kangâmiut dyke can be seen in the south-facing anorthosite wall in the far background (white arrow). b: detail of sharp contact between kangâmiut dyke (leftmost in a) and anorthosite. note subtle (primary?) layering in dyke. notebook (red dashed circle) is 11 x 18 cm. c: irregular contact between kangâmiut dyke and anorthosite host, suggesting only minor rheology contrast between dyke and host rock during emplacement or a subsequent structural overprint or both. hammer (red dashed circle) is 45 cm long. d: weakly foliated anorthosite intruded by grey, tonalitic sheet, similar to those shown in fig. 17. notebook (red dashed circle) is 11 x 18 cm. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 22 of 96 geusbulletin.org linear structures include: (1) minor fold axes; (2) lineation defined by elongate mineral aggregates (stretching lineation); (3) lineation defined by a preferred dimensional orientation of crystals with elongate habits (mineral lineation). linear structures recorded in the archaean gneisses are mainly of the first two types. orientation diagrams of linear structures from qaqqatoqaq (fig. 16) show steeply se-plunging lineation that are distinctly different from those along itillip ilua. the broad map-view distribution and orientation of the dykes in these areas are shown in fig. 16. actual field mapping of dykes has taken place along the shore of itillip ilua and on a few n–s traverses across qaqqatoqaq and south of the fjord. in the outer parts of itillip ilua, the dykes have been mapped in detail by nash (1979a, b). all other dykes are traced from aerial photographs (see the simplified summary in fig. 16). all dykes are subvertical to steeply north-dipping. in the western part of qaqqatoqaq, most dykes trend ene, whereas in the eastern part, there are distinct neand nw-trending populations (fig. 16). within the kangerlussuaq area, escher et al. (1975) found that nne-orientated dykes cut the roughly e–w-trending (85°–110°) dykes and a similar relationship was reported by mengel et al. (1996). in qaqqatoqaq, however, intersections between eneto ne-aligned dykes and the nw-trending dykes are more inconsistent, with nw-trending dykes both cutting (fig. 20a) and being cut by ene-aligned dykes (fig. 20b), leaving open the possibility of a single prolonged period of intrusion of all dykes of different orientations, as noted by escher et al. (1976a) from the inussuttusup tunua area. the dykes are generally regular in form with a straight, parallel-sided appearance. on closer inspection, however, many dykes exhibit irregularities like offsets of dykes and apophyses (fig. 20c). variation in thickness along dykes over relatively short distances has been observed in only a few dykes in these areas (fig. 20e). the dykes can be up to 60 m wide, with the mode being around 15 m. dykes are dark brown to black, rather homogeneous, and without phenocrysts. thicker dykes (>5 m) usually have finer-grained 1–20 cm wide margins, depending on the width of the dyke. towards coarser-grained centres, larger dykes grade into doleritic and rarely gabbroic textures. smaller dykes (<5 m) usually show a doleritic texture throughout except for a 2–3 cm wide fine-grained, often chilled, margin. the fine-grained margins usually have a porphyritic texture. in outcrops, where the geometry of foliations both inside and outside dyke margins can be established accurately, their interrelationship usually implies simple shear deformation of the dyke, with respect to its host. such margin-parallel shear zones may occupy one or both sides of the dyke (figs 20e, g), or in the case of wider dykes, even develop through its centre (fig. 20f). there is a clear pattern of a sinistral shear component along ne dykes and dextral shear along nw dykes (fig. 21), as described by jack (1978) and escher et al. (1976a) from the inussuttusup tunua area. the preservation of subtle, brittle features along dyke margins may suggest that dykes preferentially intruded into pre-existing zones of weakness (fractures, foliations) in the host gneisses. additionally, the fact that planar fabrics developed in and along the dykes could indicate a phase of localised deformation after the emplacement of the dykes. 3.1.6.2 itilleq shear zone along the north and south shores of itillip ilua, the n-dipping compositional layering and schistosity of the gneisses is uniformly e–w-trending, as opposed to the oblique ne to se trends across qaqqatoqaq and the e to ne trends across nuussuup qulaa (fig. 16). the c. 7 km wide zone along itillip ilua is made up of a number of e–w-trending panels of stronger deformation, which collectively constitute the itilleq shear zone (nash 1979a; fig. 16). individual panels of higher deformation a b c d e f g 100 m 50 m 40 m 10 m 2 m 2 m 20 m 82 75 84 65 80 80 87 72 85 82 69 72 80 72 75 75 80 60 68 78 75 88 65 62 80 74 56 68 40 69 89 78 8080 63 75 70 68 n fig. 20 examples of dyke geometries and relationships between dykes of different orientations in the qaqqatoqaq area (reproduced from korstgård 1980). a: wnw-trending dykes cut ne-trending dykes. b: nw-trending dyke cuts across ne-trending dykes, but also turns into an ne-trending orientation. c: ne-trending dyke with wnw-trending apophyses. d: ene-trending discordant dyke with apophyses showing right-stepping intrusion geometry. e: wnw-trending dyke shows dramatic thickness variations along strike. in thicker parts of dyke, its margin and the host gneiss show foliation development. in thinner parts of dyke, foliation is developed across the entire dyke. f: sharply discordant ene-trending dyke with internal foliation showing sinistral asymmetry. g: structures in a host gneiss that rotates asymptotically into parallelism with margin of undeformed, discordant ene-trending dyke. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 23 of 96 geusbulletin.org west itillip ilua east itillip ilua nn (n = 48) (n = 52) (n = 14)(n = 12) fig. 21 rose diagrams showing orientation of kangâmiut dykes at itillip ilua. sinistral (orange) and dextral (grey) shear along or in margins of dykes from western (left) and eastern (right) itillip ilua. in both cases, conjugate bimodal trends and shear distributions are roughly consistent with overall n–s compression. the obtuse bisectrix in western itillip ilua is orientated c. nnw, whereas in the east it is orientated n–nne, suggesting a rotation of the stress field responsible for the observed distribution of shears. note some overlaps between the two dyke sets in the acute bisectrix. source: korstgård & park (unpublished data). fig. 22 penetratively deformed tonalitic to granodioritic amphibolite facies archaean gneiss. disrupted mafic agmatite layers are visible beneath the notebook (yellow dashed circle). lateto post-kinematic granitic veins alternating with panels of straight gneiss fabrics are also shown. eqalugaarsuit, north shore of itillip ilua. notebook is 11 x 18 cm. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 24 of 96 geusbulletin.org are usually between a few metres and several tens of metres wide and alternate with less deformed areas (fig. 22). they can not typically be traced for any significant distance along the strike. the apparent horizontal sense of shear inside these high deformation zones – based on regional patterns and on bends and drags along its boundaries – appears to have been dextral. the dominant direction of movement along these vertical to steeply n-dipping planes – as estimated from the stretching lineations on schistosity planes – plunges moderately to steeply to the w and wnw and is observed in both the gneissic host rocks and the dykes (fig. 23a–c; summarised in the orientation plots of fig. 16). in apparent contradiction with those observations, there are a few spectacular sinistral delta-clasts on the horizontal surfaces on some of the high-strain gneiss panels (fig. 24) with apparent rotation axes that parallel the previously described steeply west-plunging stretching lineation. following sanderson & marchini (1984) and robin & cruden (1994), these observations are interpreted to reflect a local, but significant component of localised transpressive flow, perpendicular to the lineation. while these features are rare, they do support the presence of overall n–s transpressional shortening, similar to that described along dyke margins in section 3.1.6.1. in the high-strain panels, gneisses gain a very strong, almost mylonitic, planar fabric (fig. 25a, b). syn-kinematic leucocratic veins are transposed into parallelism with this planar fabric and often only survive as feldspar clasts (fig. 24, 25b, 26a). linear structures are equally well developed and are mostly composed of elongate quartz and feldspar aggregates (fig. 23). mafic layers (some of which could be dykes emplaced prior to shearzone formation) likewise are converted to elongated lenses or agmatitic zones (figs 22, 25c). 3.1.6.3 dykes in the itilleq shear zone the pattern of distribution and orientation of the kangâmiut dykes shows a marked change along itillip ilua where there is a concentration of dykes with an 80–90° orientation in the western part and a 90–100° orientation in the eastern part (fig. 16). this slight change in the trend of the dykes coincides geographically with a similar change in the gneissic fabrics in the e–w-trending itilleq shear zone (fig. 16). most dykes are vertical or dip steeply to the north. many of the dykes in the itilleq shear zone show little or no sign of deformation, and a b c fig. 23 field examples of linear fabrics in dykes and gneisses. a: steeply wnw-plunging lineation in the narrow part of a podded kangâmiut dyke. the dyke width changes from 40 m in the widest part of the pod to less than 1 m in the thin part connecting the pods. nuussuaq, south shore of itillip ilua. scale is 10 cm. b: steeply wnw-plunging linear aggregates of quartz and feldspar inside a strongly sheared grey amphibolite facies gneiss panel. itilleq shear zone. visible part of hammer is 30 cm. c: steep lineation in quartzo-feldspathic gneiss. qaqqatoqaq. pen is 14 cm. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 25 of 96 geusbulletin.org where neor nw-trending dykes occur with e–w dykes in the marginal zones of the itilleq shear zone, they are not less deformed (as would be expected if they represented un-sheared relics) but are similar to their e–w counterparts. this regional deflection of the dykes can only be satisfactorily explained by a primary local change in the orientation of the stress field (i.e. dominant fabric in the host gneisses), which guided the intruding dykes. such an observation was also noted in the inussuttusup tunua area (jack 1978). due to the dominance of c. e–w-dyke orientations, very few intersections were observed inside or along the a b fig. 24 field examples of sinistral shear from eqalugaarsuit on the north shore of itillip ilua a: sinistrally rotated, δ-type winged feldspar porphyroclast in high-strain gneiss panel viewed on a horizontal surface. scale is 10 cm. b: sinistrally rotated δ-type blocky feldspar porphyroclast in highly strained gneiss on horizontal surface. upper part of photo shows discordant late-kinematic pegmatite. scale is 10 cm. b a c fig. 25 field examples of gneiss fabrics from itillip ilua a: mylonitic zone in host gneiss adjacent to kangâmiut dyke (upper part of photo). nw nuussuaq, north shore of itillip ilua. pen measures 13 cm. b: wnw-plunging stretching lineation in strongly sheared mafic gneisses, with several generations of discordant pegmatite veins. this outcrop represents one of the many high-strain panels in the itilleq shear zone. sinistral δ-clast seen in fig. 24a is from this outcrop. just nw of eqalugaarsuit on the north shore of itillip ilua. scale is 10 cm. c: typical layered, polyphase, tonalitic to granodioritic amphibolite facies gneiss with mafic agmatites and several generations of pegmatites, some only mildly deformed (under hammer). this outcrop is typical of the basement gneisses outside the high-strain panels (e.g. fig. 25b). eqalugaarsuit. hammer is 45 cm. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 26 of 96 geusbulletin.org itilleq shear zone. an exception is the dyke intersections visible on the north shore of itillip ilua just north of equalugaarsuit where a large wnwto nw-trending dyke cuts several smaller neto ene-trending dykes (figs 20a, 27) in the margin of the itilleq shear zone as defined in fig. 16. both inside and outside the shear zone there are examples of dykes changing orientation (e.g. fig. 20b) making any chronological distinctions between dyke sets – based on trends – doubtful. the simplest explanation is that the dykes were emplaced during a single protracted pulse. however, until these sets of dykes are dated, this remains speculative. in contrast to the dykes on qaqqatoqaq and nuussuup qulaa, the dykes hosted by the strongly deformed gneisses in the itilleq shear zone can be highly irregular in thickness and intrusion geometry. below, we present the two main dyke types: (1) podded dykes, which are exposed as pods and lenses, often separated or connected by very thin segments (fig. 28), and (2) regular dykes with small, low angle apophyses and bayonets along their margins. we use the term ‘podded’ as a strictly non-genetic description of the dykes in question. while ‘pinch and swell’ would accurately describe the geometry, that term is typically restricted to phenomena of a tectonic origin (e.g. fossen 2016, p. 316). as some features of the shapes and intrusion forms described in this section are proposed to be of primary intrusive origin, we propose ‘podded’ as the best neutral term for the observations. one of the best examples of a podded dyke is exposed across the nuussuaq peninsula on the south shore of itillip ilua (fig. 16). this dyke was mapped along an exposed length of c. 1.5 km and is characterised by relatively long pods connected by short, thin segments (fig. 29). the shape of these pods in the horizontal section is elliptical and in the few cases where partial vertical sections were observed, these also appear elliptical. this 3d geometry is also suggested in horizontal sections by inclined dips of dyke contacts between country rock and dyke. attempts to find outcrops that exposed the entire vertical geometry of a pod have been unsuccessful. it is therefore unknown if the pods are oblate or prolate spheroids or even triaxial ellipsoids. the podded dykes are mostly concordant, although some local discordances can be found (fig. 28c). a b c fig. 26 highly strained gneiss panels in the itilleq shear zone. a: detail of one high-strain panel in tonalitic-dioritic amphibolite facies gneiss, with widespread remnants of pegmatites, preserved as porphyroclasts (detail of rotated, winged porphyroclast in fig. 24a). just west of eqalugaarsuit, north shore of itillip ilua. scale is 10 cm. b: straight, homogeneous amphibolite facies gneiss with steep stretching lineations. south shore of inner itillip ilua. hammer in circle is 40 cm. c: detail of highstrain panel in dark, tonalitic-dioritic gneiss with late-kinematic pegmatite showing subtle development of sinistral shear bands and an angular hornblendite fragment showing no rotation in this horizontal exposure. just west of eqalugaarsuit. scale is 10 cm. fig. 27 large wnw-trending dyke cuts several smaller ene-trending dykes (fig. 20a). visible in the foreground is a road leading to the dye-1 radar station (part of the distant early warning line system of radar stations active from 1958 to 1988). eqalugaarsuit, north shore of itillip ilua. ridge with light coloured gneisses in foreground is c. 300 m high. white arrows indicate dyke intersections. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 27 of 96 geusbulletin.org the shape of the podded dykes and their concordance with gneiss schistosity could suggest that the dykes were deformed into parallelism with gneiss structures and that their geometry is a boudinage, meaning a tectonic phenomenon. however, certain features of the dykes contradict this interpretation. the sketch in fig. 30b shows part of a large, podded dyke from the north shore of itillip ilua. the dyke has a maximum width of about 50 m. where the dyke narrows, intricate intrusive features are seen with small inter-connected apophyses forming an irregular network. the interconnecting apophyses are discordant to the gneiss schistosity and completely undeformed, hence the narrowing of the dyke cannot have been produced by deformation and must have been an original intrusive feature. figure 30c shows a field example of the discordance between gneiss fabrics and the delicate dyke strands at the narrow tip of the pod. furthermore, although parts of the podded dykes are deformed at some localities, they generally do not appear more (or less) deformed than their straight-sided counterparts in the itilleq shear zone and locally there are podded dykes within tens of centimetres of straightsided, concordant dykes (e.g. right of the pod in fig. 28a). the e–w-orientation of the dykes in the itilleq shear zone and the general concordance of dyke margins and gneiss fabrics are therefore most likely primary features. the strong fabric in the gneisses constitutes a closed planar discontinuity. a major effect of such a discontinuity is to lower the tensile strength and shear strength along the discontinuity surface. the schistosity planes in the itilleq shear zone will therefore be preferred sites of the fracture arrays forming ahead of dyke magma emplacement, and the geometry of the fractures will be controlled by the stress affecting the rocks. as magmatic pressure increases, the fractures will dilate into one of several final dyke forms observed along itillip ilua. nash (1979b) suggested that the podded form of many of the dykes in western of itillip ilua in 3d resembles that of an upward-pointing hand with the narrow parts connecting individual fingers (fig. 3 in nash 1979b). observations of podded dykes in the eastern part of itillip ilua, however, suggest that the 3d shape may be more like an egg carton, displaying variations of pod geometry in both horizontal and vertical directions (fig. 31). b a b c fig. 28 features in podded kangâmiut dykes. a: train of several pods (yellow outline) adjacent to un-podded thin dyke (green outline). north shore, mouth of itillip ilua. pod in foreground (marked as b) is 1 m wide. b: detail of pod shown in panel a. note stronger deformation along dyke margins and in the adjacent host gneiss. north shore, mouth of itillip ilua. notebook is 11 x 18 cm. c: pod of kangâmiut dyke showing discordance with gneissic layering (especially clear on left side near pen). towards bottom of picture, dyke thins to a few centimetres before joining another, wider pod. south shore of saqqap kangerluarsua. pen measures 13 cm. n ? 89 55 85 63 68 80 73 40 65 200 m 50 m fig. 29 sketch of podded dyke across a small peninsula at nuussuup qulaa on the south shore of itillip ilua (location in fig. 35). note how the apophysis in the enlarged area is unaffected by the drastic changes in thickness, observed just a few metres away in the large dyke. “?” is not exposed. reproduced from korstgård (1980). https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 28 of 96 geusbulletin.org the distribution and orientation of dykes around itillip ilua (including qaqqatoqaq, nuussuup qulaa and the itilleq shear zone) seem unrelated to the amount of deformation in the dykes solely based on visual observations. dykes in this area are deformed internally and generally not together with their country rocks. the extent of deformation in the dykes shows considerable variation, which does not, however, correlate with their position either outside or inside the itilleq shear zone. in the field, dyke deformation is expressed as a schistosity defined by a parallelism of flattened feldspar aggregates and amphibole aggregates. in coarser-grained dykes, an alignment of amphiboles is visible (figs 23a, c). a lineation formed by elongate aggregates of amphibole and feldspar accompanies the schistosity in many dykes and often becomes the dominant structural element, especially in the thin zones connecting the pods (figs 23a, 32). n 56 74 89 84 c b a 10 m 20 m fig. 30 intrusive features in podded dykes from the north shore of itillip ilua. a: field sketch of 20 m wide kangâmiut dyke that over 20 m narrows to a couple of centimetre-wide strands. reproduced from korstgård (1980). b: field sketch of 50 m wide kangâmiut dyke pod at the end of which is an intricate network of interconnected apophyses that are discordant to the host rock foliation and entirely undeformed. this suggests that the thinning of the pod is a primary feature rather than the result of deformation. from korstgård (1980). c: delicate discordant features preserved along thin dyke strands at tip of pod in kangâmiut dyke. hammer is 40 cm. fig. 31 interpretive sketch of the 3d geometry of podded dykes. based on field observations of mostly horizontal and a few vertical outcrops, as well as structural readings along pod margins. pods can vary in thickness from 1 m to 60 m. reproduced from korstgård (1980). fig. 32 strongly deformed tonalitic-granodioritic host gneiss adjacent to narrow part of podded dyke (upper part of figure). the dyke does not exhibit the same structures as seen in the host rocks. stretching lineation (seen as mineral aggregates in right side of image) plunge to wnw. south shore, head of itillip ilua. hammer is 40 cm. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 29 of 96 geusbulletin.org similar to dykes outside the itilleq shear zone, dykes inside the zone can show shear on one or both margins of the dyke and even through their centres. in more intensely deformed dykes, and very often in the narrow parts connecting the pods, the shearing affects the entire dyke width, and the schistosity becomes concordant or near concordant with the margins of the dyke (fig. 33). the sub-vertical to steeply wand wnw-plunging aggregate lineation in dyke margins and in thin zones between pods are interpreted to have formed during overall n–s compression with a considerable vertical component. another population of dykes display clear discordance with the sheared gneisses and preserve delicate boundary features such as thin, low angle apophyses, bayonets, and bridges (examples in fig. 34). while these dykes clearly intruded relatively brittle host rocks long after the formation of its penetrative regional fabric, an apparent and consistent asymmetry of the delicate intrusion features carries information about the geometry of the stress field encountered by the dykes. the dykes can conveniently be split into two categories based on the asymmetry of apophyses along their sidewalls, into either ‘left-stepping’ or ‘right-stepping’ intrusions (see fig. 34). our observations, both outside and inside the itilleq shear zone, indicate that dykes that trend north of east (‘counter-clockwise’ relative to the dominant gneissic fabric) are fig. 33 view from the widest part of the pod towards the narrow part (outlined in yellow dashed lines) on the south shore of inner itillip ilua. towards the narrow part, both dyke and host gneisses are increasingly strongly deformed, especially along the dyke’s left margin. late, discordant leucocratic pegmatite truncates the narrow zone. notebook is 11 x 18 cm. a b c d fig. 34 intrusion features in kangâmiut dykes in itilleq shear zone. a: strongly sheared, straight gneiss truncated by late-kinematic granitic pegmatite (clockwise relative to gneissic fabric) in turn cut by undeformed thin kangâmiut dyke (counter-clockwise). eqalugaarsuit, north shore of itillip ilua. scale is 10 cm. b: neto ene-trending kangâmiut dyke cutting host gneisses in a counter-clockwise orientation relative to gneiss layering and showing overall left-stepping intrusion geometry. preservation of delicate apophyses suggest no post-emplacement deformation. eastern nuussuaq, south shore itillip ilua. scale is 10 cm. c: 5 cm wide nne-trending kangâmiut dyke cutting heterogeneous granitoid with supracrustal xenoliths (not seen in picture) and showing right-stepping intrusion asymmetry. this outcrop is south of the itilleq shear zone. south shore, mouth of itillip ilua. coin diameter is 21 mm. d: margin of a discordant, left-stepping kangâmiut dyke (main body to lower left in picture) with a delicately preserved thin apophysis that cuts the straight planar fabric in its host gneiss (from middle of photo towards top left, shown by arrows). eqalugaarsuit, north shore itillip ilua. scale is 10 cm. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 30 of 96 geusbulletin.org typically left-stepping, whereas dykes that trend clockwise relative to the gneissic fabric are right-stepping. the stress system responsible for these features is consistent with an approximately n–s-directed maximum horizontal stress. there are possibly minor fluctuations towards nnw–sse in west itillip ilua and towards nne–ssw in the eastern part, based on the subtle variations in the orientations of both the hosting straight belts and dykes from west to east. in westernmost itillip ilua, nash (1979a) also reports detailed evidence for greatest compressive stress (σ1) being orientated approximately n–s during stage-2 shears accompanying or slightly post-dating dyke emplacement. while the two types of dykes described above are quite different in appearance, they clearly occupy the same temporal window (i.e. pre-nagssugtoqidian orogeny and post-straight belt structures in the itilleq shear zone). there are at present no conclusive chronological data to identify their relative ages. if they are broadly the same age, as we suspect, the differences in geometry are indicative of variations in host gneiss properties and the ambient stress field at the time of intrusion. the dykes truncating the straight fabrics in the itilleq shear zone intruded brittle rocks and clearly escaped subsequent deformation as shown by the preservation of subtle intrusion features. in contrast, the podded dykes and the dykes outside the itilleq shear zone, while still being discordant and displaying a plethora of primary intrusive features, did subsequently deform along their margins, presumably related to nagssugtoqidian deformation. further analysis is needed to refine this sequence of events. 3.2 dykes and country rocks of the ikertooq region (from saqqap kangerluarsua to maligiaq) 3.2.1 saqqap kangerluarsua (kangerluarsuk) the quartzo-feldspathic rocks along the southern shore of saqqap kangerluarsua (kangerluarsuk) are part of the archaean gneisses of qaqqatoqaq and nuussuup qulaa. in the western part of the south shore of saqqap fig. 35 simplified geological map of the itillip ilua – ikertooq region. the ikertooq shear zone is located along the central supracrustal bands in maligiaq and roughly follows itillip ilua. after grocott (1977), allaart & jensen (1979), jack (1978) and korstgård (1980). sisimiut amerloq ikertooq m al ig ia q akulleq avalleq qeqertalik saqqap kangerluarsua qaqqatoqaq nuussuup qulaa itillip ilua itillip ilua itilleq inussuttusup tunua sarfannguit n unaat 66°30´53°00´ 66°45´ 10 km “charnockitic” gneiss granulite facies gneiss amphibolite facies gneiss anorthosite supracrustal gneiss amphibolite dyke metadyke n https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 31 of 96 geusbulletin.org kangerluarsua (fig. 35), rocks are in granulite facies, and to the east these are in amphibolite facies. the kangâmiut dykes crosscut structures and lithologies of the quartzo-feldspathic gneisses inside both granulite and amphibolite facies areas. at several localities along the southern shore of saqqap kangerluarsua, weak nw–se-trending archaean structures (fig. 16) are transected by e–w-trending, subvertical zones of ductile deformation, akin to the pre-dyke deformation zones at itillip ilua. these zones are often also loci of intrusion of the kangâmiut dykes. pre-dyke deformation zones can be found at several localities across qaqqatoqaq but are particularly well exposed along the southern shore of saqqap kangerluarsua and characterised, as in the itilleq shear zone, by ductile deformation along e–w-trending and subvertical shear planes. the shear zones along the southern shore of saqqap kangerluarsua are discrete and wellspaced but are much more numerous than at qaqqatoqaq. farther north into saqqap kangerluarsua, such steep deformation zones become still more numerous and eventually pervade the gneisses completely. along the northern shore of saqqap kangerluarsua, these are partially overprinted by subsequent nagssugtoqidian deformation and become more difficult to recognise. in stark contrast to the south shore, the planar and steep fabric of both stronger deformations and dykes become deformed and inter-folded along the north shore of saqqap kangerluarsua and in qeqertalik, inside and during what is considered to be the nagssugtoqidian deformation proper (figs 36a, c, 37b). while kangâmiut dykes in itillip ilua, qaqqatoqaq and nuussuup qulaa and along the southern shore of saqqap kangerluarsua may be sheared either along their margins or internally, these dykes are never folded. this situation changes radically at the head of and along the northern shore of saqqap kangerluarsua. from here and northwards, dykes are always thoroughly deformed, metamorphosed and folded together with their country gneisses. this transition may be illustrated by a series of field examples, where fig. 36b shows an example of relatively weak nagssugtoqidian deformation, superimposed upon gneisses with apparently archaean structures. the dyke, though thoroughly deformed and metamorphosed, cross-cuts the gneiss structures and shows very little, if any, effects of nagssugtoqidian folding. figure 36c, from a small peninsula at the head of saqqap kangerluarsua, shows an example of nagssugtoqidian folding of both the dyke and its country gneisses. the gneisses are moderately affected by nagssugtoqidian deformation, and a c b d fig. 36 field examples of gneisses from the eastern south shore of saqqap kangerluarsua. a: folded quartzo-feldspathic gneisses. b: metadyke cross-cutting quartzo-feldspathic gneisses. compass (by red arrow) is 12 cm. c: folded metadyke in gneisses. d: fold in gneisses wrapping around metadyke. hammer is 40 cm. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 32 of 96 geusbulletin.org local discordances between dyke and gneiss are preserved. in other places with even stronger nagssugtoqidian overprints, dykes and gneisses become conformable and folded together (fig. 36d). along the northern shore of saqqap kangerluarsua, dykes and gneiss structures are mostly conformable (fig. 37a) and often folded. major folds are near isoclinal (figs 37b, c) with shallow, west-plunging axes. the general orientation of planar and linear structures in saqqap kangerluarsua is significantly different, however, from what was observed at qaqqatoqaq (fig. 16). planar structures are now striking between nne and e, while linear structures dip shallowly to moderately to the wnw (fig. 38). 3.2.2 qeqertalik and avalleq the area from qeqertalik to avalleq (fig. 35) is characterised by large-scale folded country rock fabrics and metadykes that are metamorphosed in low amphibolite facies (see section 6.2). the country rocks now also include significantly more layers of supracrustal rocks, mainly garnet-biotite gneisses, mica schists and amphibolites (figs 35, 39a–d). the mineral parageneses in the quartzo-feldspathic gneisses are quartz, alkali-feldspar, plagioclase, green biotite, epidote and subordinate muscovite and blue-green amphibole. mica schists and garnet-biotite gneisses of the supracrustal units have garnet and graphite in addition to the above-mentioned minerals, but not muscovite, epidote or amphibole. the dykes consist mostly of plagioclase and amphibole with minor amounts of titanite and quartz and are thoroughly deformed and metamorphosed. garnet is rare in the metadykes but may occur in amphibolites associated with supracrustal gneisses. in the quartzo-feldspathic gneisses, it is difficult to determine whether parageneses are a result of preor post-dyke metamorphism since the country rocks in itillip ilua and saqqap kangerluarsua – that were affected by only pre-dyke metamorphism – show the same parageneses. only in a few cases, where epidote or biotite crystallised to form part of a post-dyke structure (e.g. axial plane schistosity or mullion), can the relative age of the mineral assemblage be established. however, considering the thoroughness of the metamorphism of the kangâmiut dykes and the pervasive change in structure of the gneisses, it seems most likely that the country rocks also recrystallised after the dykes, during regional nagssugtoqidian deformation and metamorphism. the gneisses show strong linear and planar fabrics. the planar fabric is usually a combination of a fine compositional layering and a biotite schistosity (fig. 40a–c). the general trend of the planar fabric is nne–ssw in qeqertalik (fig. 38), with a moderate dip to the nnw, i.e. only a slight change from the orientations in saqqap kangerluarsua (fig. 38). the linear structures are mainly stretching lineations, defined by aggregates of quartz and feldspar and minor fold axes and to a lesser extent mineral lineation defined by preferred habit orientation of elongate minerals. the folds in the area affect the compositional layering and schistosity of gneisses (fig. 41a) as well as dykes (fig. 41b). major folds tend to be tight, with axial planes dipping steeply to the north and axes plunging gently to the west. these structures are clearly outlined + c b a 50 m n fig. 37 field examples of dykes from the north shore of inner saqqap kangerluarsua a: mostly conformable relationships between gneissic banding and metadyke orientation. hammer is 40 cm. b: tightly folded metadyke with shallow north-dipping axial plane and an axis that plunges shallowly towards the west. height of section c. 200 m. c: sketch map, including locality b. discordant pegmatites are shown in red. filled circles in the stereogram represent poles to measured foliation. open circles plot measured lineation. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 33 of 96 geusbulletin.org sisimiut amerloq avalleq akulleq qeqertalik saqqap kangerluarsua ikertooq 10km 66°45´ 53°00´ a: inner maligiaq b: outer maligiaq – akulleq c: avalleq d: qeqertalik e: saqqap kangerluarsua f: s. sarfannguit nunaat – sallersua g: amerloq – n. sarfannguit nunaat h: sisimiut (n = 35) (n = 241) (n = 34) (n = 101) (n = 53) (n = 130) (n = 98) (n = 265) (n = 62) (n = 157) (n = 375) (n = 197) (n = 87) (n = 154) (n = 116) (n = 148) planar measurements linear measurements amphibolite facies gneiss granulite facies gneiss supracrustal gneiss amphibolite dyke metadyke charnockitic gneiss m al ig ia q sarfannguit n unaat sallersua fig. 38 geological map (location in fig. 2), with stereographic plots that show the poles of planar (blue) and linear (red) measurements in the saqqap kangerluarsua – ikertooq region. these points are further contoured, using stereonet v 11 software (https://www.rickallmendinger.net/stereonet) with kamb contouring, interval 2σ. a–e: data from korstgård (1980). g–f: data from grocott (1977). h: data from davidson (1978). planar and linear structures defined in fig. 16. supra xx supra xx supra xx supra xx dc ba fig. 39 field examples of supracrustal sequences (supra xx). a: steeply nw-dipping supracrustal sequence in the south-easternmost part of qeqertalik. height of cliff c. 150 m. b: moderately nw-dipping supracrustal sequence along east shore of avalleq. height of cliff in foreground c. 20 m. c: shallowly nw-dipping supracrustal sequence along east shore of avalleq. d: steeply nw-dipping supracrustal sequence in avalleq. height of cliff c. 250 m. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ https://www.rickallmendinger.net/stereonet korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 34 of 96 geusbulletin.org c b a fig. 40 field examples of banded gneiss from inner qeqertalik. a: conformable metadyke in banded gneiss. notebook at arrow is 18 cm long. b: banded gneiss that is both folded and cut by a discordant pegmatite. hammer is 40 cm long. c: conformably folded metadyke and gneiss banding. hammer is 40 cm long. fig. 41 field examples from innermost qeqertalik. a: nw-verging folded banded gneisses looking west. person for scale. b: folded metadyke looking east. height of section c. 50 m. ba https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 35 of 96 geusbulletin.org by contrasting metadykes (fig. 41b), by supracrustal bands and by planar fabrics in the quartzo-feldspathic gneisses. the kangâmiut dykes in qeqertalik and avalleq were thoroughly metamorphosed and deformed during nagssugtoqidian orogeny. no trace of their original igneous microstructure is preserved. however, original discordances are often observed and the metadykes are still clearly recognisable as originally intrusive bodies (fig. 42a, b). planar structures in the dykes are mainly a schistosity defined by the preferred orientations of elongated amphiboles. occasionally, a compositional layering on a centimetre scale is developed, which appears to have started as a segregation of plagioclase inside the metadykes. linear structures are also defined by the preferred orientations of elongated amphiboles or appear as stretching lineation defined by aggregates of plagioclases (fig. 42c). in addition to the major folds, described here, thin metadykes often display smaller scale folds (fig. 40c), but no axial planar schistosity was found associated with these folds. boudinaged (disrupted) metadykes were found in the northern part of avalleq (fig. 42d). there is considerably less variation in the general orientation of planar and linear structures in qeqertalik and avalleq (fig. 38) compared to those in saqqap kangerluarsua. 3.2.3 akulleq and maligiaq the area between akulleq and maligiaq is characterised by strong deformation of both dykes and their country rocks, as well as an increase in metamorphic grade from amphibolite to granulite facies. metamorphic mineral assemblages in the rocks change radically along outer avalleq–akulleq. in the quartzo-feldspathic gneisses epidote and muscovite disappear and give rise to assemblages characterised by biotite and amphibole, in addition to feldspar and quartz. further north, in maligiaq, metamorphic clinopyroxene and orthopyroxene appear simultaneously with a decrease or near disappearance of alkali-feldspar in the quartzo-feldspathic gneisses. in the metadykes between the southern shore of akulleq and midway into maligiaq, clinopyroxene, and to a lesser extent garnet, become important members of the mineral assemblage (fig. 35). farther north, in the inner parts of maligiaq, orthopyroxene appears. continuing across akulleq, from avalleq, and especially towards the outer parts of maligiaq, there is an increase in both the intensity and parallelism of the planar fabrics of gneisses (fig. 38). the intensity of this planar fabric culminates in the central part of maligiaq (fig. 38), where the planar fabric strikes e–w and dips around 60° towards the north (fig. 43a). the planar fabric here developed as a combination of a fine compositional layering and a biotite schistosity (fig. 43c). along ikertooq and in the outer parts of maligiaq, linear structures (stretching lineation and mineral lineation) plunge shallowly to the west. about midway into maligiaq, however, there is a sudden change in the direction of the linear structures from shallowly west to steeply north-plunging. this change in direction coincides with the beginning of an intensification of planar fabrics. the change in direction of the linear structures also approximately coincides with the onset of granulite facies metamorphism. however,  from field observations it is clear that the beginning of granulite facies is well within the zone of strong deformation, about midway into maligiaq. on a map scale (fig. 35), the planar structures along akulleq and in maligiaq define a linear zone that runs approximately east to west. this linear zone can be followed towards the west on sarfannguit nunaat (sarfannguaqland; grocott 1977) and eastwards on aerial photographs for at least an additional 50 km. the linear nature of the zone is emphasised by supracrustal bands (fig. 43a). the intensity of deformation decreases in the northern part of maligiaq, where e–w panels of low strain (fig. 43b) alternate with zones of strong deformation (fig. 43c). further north into maligiaq, nagssugtoqidian deformation appears to be quite low, as judged by a decreased intensity of fabrics and tightness of folds in dykes and gneisses. dykes and compositional layering of gneisses (fig. 43d) are only gently folded, in contrast to the tight folds and strong schistosity characterising the gneisses in the higher strain areas. based on the evidence summarised here and from previous contributions (noe-nygaard & ramberg 1961b; watterson 1974; grocott 1977, 1979; korstgård 1979a, b), the nagssugtoqidian deformation along and north of ikertooq is interpreted as a zone of strong deformation, characterised by ductile overthrusting towards the sse along a moderately dipping thrust plane. the estimated width of the overthrusting zone in maligiaq is about 5 km and its southern boundary is located about 2.5 km north of the maligiaq mouth. the amphibolite-granulite facies boundary is located about 4 km north of the mouth and appears to parallel the thrust zone in the horizontal as well as in the vertical plane. the metadykes in maligiaq occur in three different metamorphic-structural settings relative to the thrust zone defined here: https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 36 of 96 geusbulletin.org dc ba fig. 42 field examples from north-east shore of avalleq. a: banded gneiss with older mafic bodies conformably cut by c. 1 m wide kangâmiut metadykes. b: conformable metadyke in gneiss. c: stretching lineation in metadyke defined by plagioclase aggregates. hammer is 40 cm. d: disrupted metadyke in gneiss. large pod is c. 10 m long. fig. 43 field examples from around maligiaq. a: steeply north-dipping and orange weathering supracrustal sequences, where gneisses to the right in the picture contain conformable metadykes. height of main cliff about 400 m. near mouth of maligiaq. b & c: banded gneisses and metadykes folded inside a low deformation zone in inner maligiaq. dyke at water’s edge in b is approximately 1.5 m wide. d: thin metadyke conformably folded in a high deformation zone in maligiaq. hammer is 40 cm. dc ba https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 37 of 96 geusbulletin.org 1. along the north shore of ikertooq, metadykes are mostly conformable to gneiss structures and lithological boundaries (fig. 44a), but locally, discordant relationships to the country rocks are preserved (fig. 44b). the metadykes have a strong planar fabric, defined mainly by preferred orientation of elongated amphibole aggregates, and commonly a well-developed stretching lineation that plunges shallowly to the wnw (fig. 38). mostly tight to isoclinally folded metadykes are common (fig. 44c), with fold axes that are parallel to the lineation. 2. within the zone of strong nagssugtoqidian deformation (i.e. the thrust zone, immediately north of maligiaq), all dykes are parallel to the planar fig. 44 dyke-gneiss relationships in northern ikertooq. a: steeply n-dipping metadyke is broadly conformable with gneiss banding. minor discordances, however, are preserved in the foreground. b: tightly folded and steeply n-dipping dyke with rare discordant relationships with its host gneiss. hammer is 40 cm. c: subhorizontal hinge zone of synformally folded metadyke, enhanced by felsic veins. hammer is 40 cm. a, b and c photos from north shore of ikertooq close to mouth of maligiaq, looking west. d: disrupted metadykes in a 50 m wide, steeply n-dipping high-strain zone in amphibolite facies gneisses. east shore, mouth of maligiaq. dc ba https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 38 of 96 geusbulletin.org fig. 45 field examples from inner maligiaq. a: discordant and little deformed metadyke. hammer is 40 cm. b: metadyke that is conformably z-folded in granulite facies gneisses. hammer is 40 cm. ba structures in the gneisses and mostly straight and parallel-sided (fig. 44a). however, towards the south, within the amphibolite-facies part of the thrust zone, dykes may be broken up or boudinaged (fig. 44d). 3. north of the thrust zone, where all rocks are in granulite facies, the metadykes are often more irregular, yet may cross-out gneiss structures (fig. 45a) and may also be folded together with the gneisses (fig. 45b). https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 39 of 96 geusbulletin.org 4 igneous mineralogy and geochemistry of the kangâmiut dykes for this study, we have compiled a suite of published and unpublished geochemical data of kangâmiut dykes from maniitsoq to ikertooq (fig. 46; appendix 1). the published studies describe the chemistry of the dykes and proposed models for their petrogenesis. however, our focus here is on identifying any regional compositional variations from maniitsoq to ikertooq and the control these could have on the post-emplacement metamorphic overprints on the dykes, if any. we consider such an analysis important and necessary as it will allow distinction to be made between mineral compositional variations resulting from varying metamorphic conditions versus those reflecting bulk rock compositional effects. rather than view these data altogether or by source, we view them geographically in 11 areas (areas 01 to 11 in fig. 46) from south to north. these fig. 46 simplified geological map of the study area, showing the distribution and sources of analysed samples and the subdivision into areas based on sample density, host rock character, structural variations and metamorphic overprints observed in the dykes. nf: nagssugtoqidian front (see also fig. 3). data sources (see supplementary file s1 and appendix 1 for details): beckmann (jack 1978); bridgwater (coll. 1972); cadman et al. (2001); hansen (1989); jack (1978); korstgård (1980); mayborn (2000); windley (1970). n f beckman (jack 1978) bridgwater (coll. 1972) cadman et al. (2001) hansen (1989) jack (1978) korstgård (1980) and korstgård et al. (this study) mayborn (2000) windley (1970) data sources: sisimiut kangaamiut maniitsoq area 01area 01 area 02area 02 area 03area 03 area 04area 04 area 05area 05 area 06area 06 area 07area 07 area 08area 08 area 09area 09 area 10area 10area 11area 11 ikertooq itillip ilua iso rto q alanngua kangerlussuatsia q kangaamiut sermiat inussuttusup tunua kangerluarsussuaq sermersuut simiutaq kangerl ussu aq 52° 67° 67° 66° 66° 54° 54° 53° 53° 50 km https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 40 of 96 geusbulletin.org areas are defined according to sample density, host rock character, structural variations and metamorphic grades of host rocks (and dykes). in using basic dykes as monitors of post-emplacement metamorphism, the exact age of the dykes is not critical. if the basic dykes are of similar composition and were not metamorphosed prior to nagssugtoqidian orogeny, they are all ideally suited for the purpose outlined here, and in chapters 1 and 2. 4.1 igneous mineralogy dykes with preserved igneous mineralogy, or with only very minor secondary mineral growth, can be found along almost the entire area from maniitsoq to itillip ilua. as we discuss in chapter 5, the degree of metamorphic overprinting on the dykes is extremely varied, both within an outcrop and across the study area, and preserved primary igneous mineralogy and texture can be found throughout. the typical primary mineral phases are plagioclase, clinopyroxene, fe-oxides and accessory minerals such as apatite, zircon and quartz. olivine was not found. primary amphibole or orthopyroxene or both are less common but do occur across the area. in the inussuttusup tunua area, jack (1978) describes how the dominant pyroxene in the e–w dykes is pigeonite, which is commonly inverted to orthopyroxene and typically intergrown with clinopyroxene. the older, black feldspar dykes (also e–w-trending) are characterised by having primary orthopyroxene (not converted from pigeonite) as the only pyroxene. the description here focuses on the main kangâmiut dyke swarm, with its broad distribution and apparent homogeneity, while other dyke sets will be mentioned when relevant. considering the large areal extent of the swarm, the dykes are fairly homogeneous. grain sizes vary from 0.1–0.5 mm in dyke margins to 1–3 mm in centres. dyke margins generally exhibit ophitic to sub-ophitic textures (fig. 47), often with large phenocrysts of plagioclase, clinopyroxene and locally what appears to be amphibole, set in a fine-grained, chilled matrix. plagioclase laths have length:width ratios of 10:1 to 20:1, and are randomly orientated, except near the dyke wall (jack 1978) where they may show alignment parallel with the contact. plagioclase composition varies from an60 to an65 (i.e. labradorite) in the core of phenocrysts to an40 (i.e. andesine) in the rim. orthoclase content is typically <2%. individual crystals of plagioclase are often clouded with reddish to brownish very fine-grained particles (figs 47b, c). these areas tend to clear up during incipient metamorphic overprints so the cloudiness may be a pre-metamorphic feature. compositionally, the clinopyroxenes fall in the augite and diopside fields in the pyroxene quadrilateral (fig. 48c). the cluster straddling the diopside-augite boundary represents analyses from clinopyroxenes in areas 08-11, including clearly metamorphic minerals from areas 10 and 11. the spread in the augite-field is dominantly from the southern areas, where igneous pyroxenes show variable post-emplacement effects, including incipient hydration (fig. 48c). orthopyroxenes have been observed and analysed in 15 samples. they show a range in compositions from en40 to en70 (fig. 48c), with the clearly metamorphic orthopyroxenes from the northernmost areas in the narrower range of en50–en60. orthopyroxenes do not show a preference for either the n–ne-trending dykes or the (rarer) e–w-trending dykes. primary amphibole is not common and potential candidates were only found in one sample from the southern part of the study area but has been reported in several previous studies. the sample in question has experienced no or only minor hydration of primary clinopyroxenes, and the amphiboles are quite distinct and appear to have crystallised at the same time as clinopyroxenes (fig. 47a). where the replacement of pyroxenes has progressed. however, similar amphiboles have not been identified. kalsbeek et al. (1978) describe dykes between maniitsoq and kangerlussuaq with brown-green amphibole occurring as a late magmatic mineral. they also state that magmatic amphibole makes up 20–40% by volume in the groundmass of the chilled margins and that further towards the centre it varies between 5 and 10% and has grown on augite and fe oxides. hansen (1989) also reported up to 10% primary amphibole in dykes in areas 02–06 (fig. 46). it is our interpretation, that most of these amphiboles are in fact of metamorphic origin. amphiboles are found in fine-grained and coarse-grained dykes. it is our view that only rare, dark, brown-green amphiboles occurring as isolated grains surrounded by plagioclase or showing poikilitic texture with enclosed plagioclases (fig. 47a), are of primary (late) magmatic origin. this is especially true where they occur in the same section as obviously secondary, blue-green amphiboles rimming clinopyroxene or opaque grains or both. this agrees with mayborn (2000), who described late-stage magmatic amphiboles overgrowing both clinopyroxene and plagioclase in a sample from a chilled dyke margin. where metamorphic amphiboles occur both in the fine-grained matrix and as rims on clinopyroxenes, the close optical similarity between the two types strongly suggests that both were formed by sub-solidus reactions. the gradual increase in the modal amount of amphibole when progressing from slightly to completely metamorphosed dykes is in accordance with this interpretation. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 41 of 96 geusbulletin.org compositionally, the primary amphiboles straddle the boundary between mg-hornblende and pargasite (following the classification scheme of hawthorne et al. 2012). figure 48a shows that igneous amphiboles display a limited, but non-unique, range relative to the remainder of the dominantly metamorphic, amphiboles analysed for this study. in dyke margins, larger mafic phenocrysts in the finer-grained matrix are often poikilitic, completely enclosing plagioclase laths. figure 47a shows examples of both amphibole and clinopyroxene with poikilitic textures, clearly showing the sequence of crystallisation in this sample. it is also notable that the few primary amphiboles appear to include more (and smaller) plagioclase grains than observed in clinopyroxenes, supporting their late-magmatic nature. representative modal analyses based on data from a range of sources are shown in fig. 49. all data sets marked ‘h’ in the figure depict the margin-to-centre variations in individual dykes between maniitsoq and itillip ilua, whereas the remainder shows the range of values from individual dykes. the main igneous minerals show a subtle trend of decreasing volumes from south to north, especially in the northern areas, where dykes are thoroughly metamorphosed and recrystallised in amphibolite and granulite facies. in general, however, the within-dyke variations can be as significant as differences between adjacent dykes and even areas. a unique feature of the thicker kangâmiut dykes is the development of symmetric compositional zoning from margin to centre, especially the late-stage leucocratic centres (fig. 50), which are interpreted to represent late-stage mobilisates or very late-stage igneous fractionates that locally intrude earlier pulses of the dyke. in a detailed study of these leucocratic centres, mayborn et al. (2008) argued that despite their andesitic composition, the centres were not the product of contamination or fractional crystallisation, but fig. 47 photomicrographs from kangâmiut dykes with no or only minor metamorphic overprints. a: chill of 30 m wide nne-trending dyke. primary clinopyroxene (left) and amphibole (right, brown-green) overgrowing a finer-grained matrix of plagioclase, clinopyroxene and ilmenite. large primary plagioclase in upper part of picture. sample 430267. south shore, mouth of kangerlussuatsiaq. b: sample 4 m from contact of 40 m nne-trending dyke. coarse-grained clinopyroxenes and subhedral plagioclases, with minor replacement of clinopyroxene by pale-green amphibole. sample bh l-186197b. 10 km nne of maniitsoq, collected by hansen (1989). c: fresh, coarse-grained clinopyroxene–orthopyroxene–plagioclase assemblage from the centre of e–w dyke. compositional zoning in plagioclases enhanced by variable ‘dusting’ by fine-grained particles. sample jk 1332, south shore, outer itillip ilua. d: fine-grained matrix of elongate plagioclase laths and interstitial clinopyroxenes. incipient formation of pale-green amphiboles on margins of ilmenite and clinopyroxene. sample jk 1328, south shore inner itillip ilua. mineral abbreviations follow whitney & evans (2010). amp: amphibole. cpx: clinopyroxene. pl: plagioclase. opx: orthopyroxene. 0.5 mm a b dc 430267 bh l-186197b jk 1328jk 1332 amp cpx cpx pl cpx opx pl pl 0.5 mm 0.5 mm 1 mm https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 42 of 96 geusbulletin.org rather the result of mixing of the evolved kangâmiut dyke parental magma with partial melts from iherzolitic phases of the underlying mantle. the mineralogy in these veins includes amphibole, garnet, albite, epidote and biotite. it has been discussed whether these assemblages should be considered igneous and linked with later stages in the emplacement process, or whether a (much) later metamorphic overprint is responsible. windley (1970) described similar zonation in wide dykes from the maniitsoq region and proposed that the initial phases of magma were sufficiently rich in water to give rise to amphibole-quartz-dolerites, and that “by the time the last magma was intruded, there was sufficient stress and fluid pressure for foliated garnet amphibolites to crystallize in the central parts of the dykes”. other authors, including jack (1978), also favoured an ‘autometamorphic’ origin of the mineral assemblages in the late-stage felsic components of the dykes. we propose a much later metamorphic origin of the garnet-amphibole-bearing assemblages (see discussion in chapter 5). 4.2 geochemistry of kangâmiut dykes in the sno and its southern foreland in this section we consider whether any systematic compositional variations exist among the dykes from maniitsoq to ikertooq. the discussion is based on a compilation fig. 48 compositions of igneous and metamorphic minerals in kangâmiut dykes (analyses not available from dykes in areas 01 and 05). a: compositions of all amphiboles analysed for this study. axes as defined in classification scheme of hawthorne et al. (2012). field of interpreted late-stage magmatic amphiboles (e.g. amphibole in fig. 47a) is grey and outlined by a dashed black line. amphibole end-member abbreviations: ed: edenite; pg: pargasite; sa: sadanagaite, tr: tremolite, mg-hbl: magnesio-hornblende, ts: tschermakite. b: compositions of all plagioclases analysed for this study. each dot represents an analysis. binsize along the y-axis is 2% an. fresh, unaltered plagioclases fall in the an50–an65 range. metamorphic plagioclases, initially in rims adjacent to amphiboles or garnets or both, show lower an contents (<an30) in areas 02–06. with higher metamorphic grades (areas 08–11), an increases and converge around an40–an50 (cf. fig. 82). c: compositions of selected pyroxenes analysed for this study. igneous clinopyroxenes (including those that have experienced initial replacement by amphibole) occupy the aug field, whereas metamorphic clinopyroxenes (areas 10–11) straddle the di–aug field. orthopyroxenes (range from en70 to en40 with clearly metamorphic orthopyroxenes showing a narrow range of en60–50. cpx: clinopyroxene. opx: orthopyroxene. pyroxene end-member abbreviations: di: diopside, hd: hedenbergite, aug: augite, en: enstatite, fs: ferrosilite. a b c area 02 area 03 area 04 area 06 area 07 area 08 area 09 area 10 area 11 20 40 60 80 20 40 60 80 camg mg fe cpx opx cafe di fsen aug hd 0 20 40 60 % a no rt hi te igneous plagioclase from ‘freshest’ dykes 0.00 0.25 0.50 0.75 1.00 0.0 0.5 1.0 1.5 2.0 c* sum: c(al + fe3 + mn3 + cr + 2ti) a* s um : a (n a + k + 2c a) area 02 area 03 area 04 area 06 area 07 area 08 area 09 area 10 ed tr mg-hbl ts sapg field of igneous amphiboles area 11 https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 43 of 96 geusbulletin.org of kangâmiut dyke chemical analyses, including major and trace elements, and selected rare-earth element (ree) data for >400 dykes where location information is available (see fig. 46 for locations and sources of data). in areas 01 to 08, the metamorphic overprints on the dykes vary from only minor hydration to strong recrystallisation, but igneous features are generally recognisable (see chapter 5). this is not so in area 10, where all dykes, along with their country rocks, are thoroughly recrystallised in amphibolite facies, and in area 11, where dykes and their hosts are in granulite facies throughout. area 09 represents a transitional zone between areas 08 and 10. while our aim is to characterise any primary magmatic compositional variations, we acknowledge that penetrative amphibolite and granulite facies metamorphism is typically associated with enhanced mobility of certain elements (e.g. bridgwater 1979; rollinson 1993), so we fig. 48 (continued) compositions of igneous and metamorphic minerals in kangâmiut dykes. a: compositions of all amphiboles analysed for this study. axes as defined in classification scheme of hawthorne et al. (2012). field of interpreted late-stage magmatic amphiboles (e.g. amphibole in fig. 47a) is grey and outlined by a dashed black line. amphibole end-member abbreviations: ed: edenite; pg: pargasite; sa: sadanagaite, tr: tremolite, mg-hbl: magnesio-hornblende, ts: tschermakite. b: compositions of all plagioclases analysed for this study. each dot represents an analysis. binsize along the y-axis is 2% an. fresh, unaltered plagioclases fall in the an50–an65 range. metamorphic plagioclases, initially in rims adjacent to amphiboles or garnets or both, show lower an contents (<an30) in areas 02–06. with higher metamorphic grades (areas 08–11), an increases and converge around an40–an50 (cf. fig. 82). c: compositions of selected pyroxenes analysed for this study. igneous clinopyroxenes (including those that have experienced initial replacement by amphibole) occupy the aug field, whereas metamorphic clinopyroxenes (areas 10–11) straddle the di–aug field. orthopyroxenes (range from en70 to en40 with clearly metamorphic orthopyroxenes showing a narrow range of en60–50. cpx: clinopyroxene. opx: orthopyroxene. pyroxene end-member abbreviations: di: diopside, hd: hedenbergite, aug: augite, en: enstatite, fs: ferrosilite. a b c area 02 area 03 area 04 area 06 area 07 area 08 area 09 area 10 area 11 20 40 60 80 20 40 60 80 camg mg fe cpx opx cafe di fsen aug hd 0 20 40 60 % a no rt hi te igneous plagioclase from ‘freshest’ dykes 0.00 0.25 0.50 0.75 1.00 0.0 0.5 1.0 1.5 2.0 c* sum: c(al + fe3 + mn3 + cr + 2ti) a* s um : a (n a + k + 2c a) area 02 area 03 area 04 area 06 area 07 area 08 area 09 area 10 ed tr mg-hbl ts sapg field of igneous amphiboles area 11 fig. 49 modal mineralogy of kangâmiut dykes from maniitsoq to ikertooq (based on point counting of thin sections). data sources: h: hansen (1989). w: windley (1970); j: jack (1978). k: korstgård (1980). all hansen (1989) data show changes from margin to centre for individual dykes from left to right. windley’s (1970) data show changes from margin to centre for several zoned dykes (1st to 4th columns: subophitic chilled margin; 5th to 6th columns: foliated amphibole-rich dolerite; 7th to 8th columns: garnet amphibolite; 9th to 10th columns: strongly foliated amphibolite). all mineral abbreviations are from whitney & evans (2010). jack (1978) data show pigeonite as “clinopyroxene”. in areas 02 to 07, “hornblende” refers to amphiboles deemed igneous by the respective authors, and “amphibole” refers to secondary or metamorphic minerals. in areas 09–11, all minerals are metamorphic. black bars indicate dyke orientations (w, j and k data are average orientations). numbers show width of individual dykes. m -> c: margin to centre variations. opq: opaque minerals. ttn: titanite. bt: biotite. grt: garnet. amp: amphibole. hbl: hornblende. opx: orthopyroxene. cpx: clinopyroxene. qtz: quartz. pl: plagioclase. hhhhhhhhhh kkkw hj opq ttn bt grt amp hbl opx cpx qtz pl 0% 10% 20% 30% 40% 50% 60% 70% 80% 90% 100% area 02 area 03 area 04 area 05 area 06 07 area 09 10 area 11 40 m 65 m 60 m 40 m 60 m 60 m 30 m (metamorphosed dykes)50 m 100 m 45 m 120 m e-w nem -> c https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 44 of 96 geusbulletin.org expected to see these phenomena in areas 10 and 11. we include these areas in our discussion, so that any patterns, trends or signatures which could be related to enhanced element mobility, may be identified. such observations are discussed further in chapter 7. the main geochemical features of the kangâmiut dykes were described by windley (1970), jack (1978), korstgård (1980), zeck & kalsbeek (1981) and bridgwater et al. (1995). in addition, petrogenetic studies were presented by mayborn (2000), cadman et al. (2001), mayborn & lesher (2004, 2006) and mayborn et al. (2008). the compositional variation shown by the dykes in the present data set clearly shows them to be sub-alkaline tholeiites (fig. 51), characterised by an overall fe enrichment pattern. the higher sio2 values (basaltic andesites: 52–56%; andesites: >56%) represent the leucocratic centres of larger dykes described previously. despite the compositional differences, mayborn et al. (2008) concluded that the late-stage andesitic components of the dykes were the products of mixing evolved kangâmiut dyke magma with small amounts of partial mantle melts (see also section 4.1). based on mixing models involving nd and sr isotopes, mayborn et al. (2008) show that the kangâmiut precursor magma experienced only minor (≤8%) crustal contamination during differentiation at depth. similar results were reported by cadman et al. (2001), who additionally suggested that while the chemical evolution of the magma was controlled by clinopyroxene-plagioclase fractionation, later stage involvement of amphibole as a fractionating phase might have been important locally. in contrast, mayborn & lesher (2006) showed that amphibole was likely not a primary fractionating phase, and instead described the amphibole occurrences similar to the one shown in fig. 47a, as a “late-stage magmatic phase”. both studies emphasised, based on ree considerations, the lack of firm evidence for garnet being part of the fractionating assemblage. molecular proportion diagrams (i.e. pearce diagrams; pearce 1968, 1970) have been criticised for potentially yielding spurious results when used to test predictions for the identity of fractionating phases of an evolving magma (e.g. rollinson 1993). none-the-less we feel that their usage here, to underscore the homogeneity of the present data set, is justified. figure 52 shows that similar, if not identical, assemblages of fractionating phases, characterise the dyke data from across a wide study area and are thereby consistent with a common differentiation process and parent for all dykes. the figure also includes data from orthopyroxene-bearing noritic dykes from sw greenland (bridgwater et al. 1995), which display distinctly different trends than the kangâmiut dykes. in cross-plots of mgo vs. other major oxides (fig. 53), it is also clear that clinopyroxene and plagioclase are the dominant precipitating phases. at higher mgo values, al2o3 broadly decreases with decreasing mgo suggesting that both clinopyroxene and plagioclase fractionated. meanwhile, the scattered increase in al2o3 at lower mgo fig. 50 leucocratic centres in larger, zoned dykes. a: felsic material in centre of a >20 m kangâmiut dyke (ene-trending) showing both intrusive relationship (front, left) and concordance (centre, right) with the main, mafic part of the dyke. north shore, mouth of itillip ilua (pen in red dashed circle measures 13 cm). b: centre of a 10 m wide kangâmiut dyke (e–w-trending) showing an irregular layer of leucocratic, coarsegrained material. late-intrusive nature of the felsic material is seen by the discordant and concordant relationships with the main mafic dyke body, and by the penetrative foliation overprinting both lithologies. eqalugaarsuit, north shore itillip ilua. notebook is 13 x 18 cm. a b https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 45 of 96 geusbulletin.org values (<3% mgo) suggests that plagioclase was less important at this stage. the strong positive correlation between mgo and cao indicates that both plagioclase and clinopyroxene fractionated. in addition, we observe decreasing values in tio2 and feo(tot) with decreasing mgo below 4.5–5% mgo, suggesting that ilmenite enters the precipitating assemblage around those values, as noted by mayborn & lesher (2006). the only elements that may show scatter attributable to mobility during metamorphism are k and na (see also section 7.1). the subtle trends seen in mgo vs. k2o are further discussed in section 7.1. a few of the outliers seen in fig. 52 are also identified in fig. 54, which shows mgo vs. cao/al2o3. while none of these elements are expected to show elevated mobility during metamorphism, the cross-plot may isolate dykes of different initial compositions. the overwhelming majority of analyses plot along a trend with fractionation of plagioclase and clinopyroxene (mayborn 2000). however, along with the scatter, there are a few dykes with higher mgo or lower cao/al2o3 values that stand out. by isolating data from each of the 11 areas, we find that the data in fig. 54 cluster around two dominant trends (determined visually). with decreasing mgo, data points follow the broad trend 1 until mgo reaches 4–5% and then shifts to the slightly steeper and better-defined trend 2 (fig. 54). this happens to coincide with the trends in both the mgo vs. tio2 and mgo vs. feo(tot) plots (fig. 53), where these shifts were interpreted to indicate that ilmenite became a significant part of the fractionating assemblage. the few data points falling below trend 1 at mgo >5% deserve some attention. five orthopyroxene-bearing, e–w-trending dykes from mayborn (2000) come from area 05 and 06 and plot at slightly higher mgo and lower cao/al2o3 values than the main cluster. the remaining outliers are from the northern areas (area 08–11) from e–w-trending dykes also with higher mgo content. these may have carried primary orthopyroxene, but due to later metamorphic overprints, none are preserved. it is worth reiterating, that while figs 51–54 reflect some compositional variations in the data set, there is no evidence that any of the dykes initially had noritic 0 1 2 3 4 0 2 4 6 feo(tot)/mgo ti o 2 zr (ppm) 100 200 300 400 500 tholeiitic (skaergaard)-tr end calc-alkaline trend 0.00 0.25 0.50 0.75 1.00 0 100 200 300 400 500 zr p 2o 5 alkaline tholeiitic 0 2 4 6 45 50 55 60 sio2 n a 2o +k 2o alkaline (basalt) sub-alkaline (basaltic andesite) (andesite) 20 40 60 80 100 20 40 60 80 100 20 40 60 80 10 0 f a m area 01 area 02 area 03 area 04 area 05 area 06 area 07 area 08 area 09 area 10 area 11 tholeiitic series calc-alkaline series a b c d fig. 51 bulk-rock compositional variations and classification of kangâmiut dykes. a: feo(tot)/mgo vs. tio2 (wt%). composition of kangâmiut dykes compared to typical tholeiitic and calc-alkaline trends (from miyashiro 1974). the differentiation recorded by increasing xand y-axis values is also shown by concomitant increases in zr (ppm; as indicated by circle size). b: zr (ppm) vs. p2o5 (wt%; after winchester & floyd 1977). a clear tholeiitic trend is shown. c: sio2 (wt%) vs. k2o+na2o (wt%). most kangâmiut dykes have low alkali content and sio2 between 48% and 55% and so plot within the sub-alkaline fields of basalts and basaltic andesites. most of the data points with higher sio2 values extending into andesites, stem from leucocratic centres of wider dykes (see text for discussion). these higher sio2 values also correspond to the higher feo(tot)/mgo values (>4.5; panel a) and to the higher zr values (>250 ppm; panel b). boundaries and fields from irvine & baragar (1971). d: kangâmiut dykes show the fe-enrichment trend typical of the tholeiitic series of basic rocks. solid dividing line from irvine & baragar (1971), stippled line from kuno (1968). a: k2o+na2o, f: feo(tot), m: mgo, all wt%. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 46 of 96 geusbulletin.org compositions as seen in many dykes in sw greenland (e.g. bridgwater et al. 1995). as mentioned previously, some of the wider kangâmiut dykes display a visible mineralogical zoning from margin to centre, with finer-grained, sub-ophitic margins and coarser-grained, often leucocratic, centres. to assess the bulk chemical variations associated with this zoning, we have assembled a collection of dykes 0.0 0.1 0.2 0.3 0.4 0.5 0.0 0.5 1.0 1.5 2.0 si/zr c a/ zr 0.00 0.25 0.50 0.75 1.00 0.0 0.5 1.0 1.5 2.0 si/zr fe m g/ zr ba opx-bearing norites from sw greenland opx-bearing norites from sw greenland opx pl (amp) cpx ol (amp) pl cpx opx, (ol) area 01 area 02 area 03 area 04 area 05 area 06 area 07 area 08 area 09 area 10 area 11 fig. 52 compositional variations and fractionation trends of kangâmiut dykes. a: pearce-type diagram (si/zr vs. ca/zr) showing bulk compositional effects of extracting precipitating phases. vectors of orthopyroxene, olivine, plagioclase, amphibole and clinopyroxene are based on average compositions of these phases in this study. trend is clearly consistent with compositional variations caused by a constant precipitating mineral assemblage (mainly plagioclase, clinopyroxene), throughout differentiation. calculation of variables follows pearce (1968). b: pearce-type diagram for si/zr vs. femg/zr, where femg is the sum of the molecular proportions of mgo and feo(tot), following pearce (1968). as in panel a, the data display a consistent trend. grey shading: data from orthopyroxene-bearing norites in sw greenland (bridgwater et al. 1995), which clearly show a distinctly different trend. opx: orthopyroxene. ol: olivine. pl: plagioclase. amp: amphibole. cpx: clinopyroxene. 45.0 47.5 50.0 52.5 55.0 10.0 12.5 15.0 17.5 20.0 0 5 10 15 0.0 0.5 1.0 1.5 2.0 0 1 2 3 4 5 0 5 10 15 20 0 1 2 3 4 5 0.00 0.25 0.50 0.75 1.00 0.0 2.5 5.0 7.5 10.0 0.0 2.5 5.0 7.5 10.0 0.0 2.5 5.0 7.5 10.0 0.0 2.5 5.0 7.5 10.0 0.0 2.5 5.0 7.5 10.0 0.0 2.5 5.0 7.5 10.0 0.0 2.5 5.0 7.5 10.0 mgo 0.0 2.5 5.0 7.5 10.0 mgo si o 2 ti o 2 a l 2o 3 fe o (t ot ) c ao n a 2o k 2o p 2o 5 area 01 area 02 area 03 area 04 area 05 area 06 area 07 area 08 area 09 area 10 area 11 fig. 53 mgo vs. major oxides for the kangâmiut dykes (all wt%). the distinct trends seen in mgo vs. tio2 and mgo vs. feo(tot) suggest that ilmenite enters the precipitating phase assemblage around mgo = 4.5%. note how alkalis (k2o, na2o) and p2o5 behave as incompatible elements, whereas cao and al2o3 show the most distinct depletion with decreasing mgo. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 47 of 96 geusbulletin.org from which margin to centre sampling traverses were available (fig. 55). each block of data (4 to 9 samples) represents a traverse from margin (left) to centre (right) of a wide (>20 m) dyke. several show subtle decreases in magnesium number (mg#) and increases in zr, from dyke margins to centres, in accordance with the fractionation trend previously described for the swarm as a whole. some dykes, however, display little variation across a dyke, while a few even show a compositional asymmetry, with lower mg# and higher zr away from the dyke centre. the overall compositional homogeneity and predictability of the data set are also shown in a multielement diagram (fig. 56a). in this figure the trace element concentrations are normalised to primitive mantle values (from mcdonough & sun 1995) and arranged from left to right in order of increasing compatibility in a small fraction melt of the mantle (e.g. thompson 1982; rollinson 1993). for simplicity, only mean values from each area are shown on a background of the distribution of results from all samples. the number of samples per area ranges from 5 (area 01) to 62 (area 6) for a total of 239 samples. the gentle, downward trend from the traditionally more mobile and incompatible elements in fig. 56a (left) and towards the increasing compatibility of fig. 53 (continued) mgo vs. major oxides for the kangâmiut dykes (all wt%). the distinct trends seen in mgo vs. tio2 and mgo vs. feo(tot) suggest that ilmenite enters the precipitating phase assemblage around mgo = 4.5%. note how alkalis (k2o, na2o) and p2o5 behave as incompatible elements, whereas cao and al2o3 show the most distinct depletion with decreasing mgo. 45.0 47.5 50.0 52.5 55.0 10.0 12.5 15.0 17.5 20.0 0 5 10 15 0.0 0.5 1.0 1.5 2.0 0 1 2 3 4 5 0 5 10 15 20 0 1 2 3 4 5 0.00 0.25 0.50 0.75 1.00 0.0 2.5 5.0 7.5 10.0 0.0 2.5 5.0 7.5 10.0 0.0 2.5 5.0 7.5 10.0 0.0 2.5 5.0 7.5 10.0 0.0 2.5 5.0 7.5 10.0 0.0 2.5 5.0 7.5 10.0 0.0 2.5 5.0 7.5 10.0 mgo 0.0 2.5 5.0 7.5 10.0 mgo si o 2 ti o 2 a l 2o 3 fe o (t ot ) c ao n a 2o k 2o p 2o 5 area 01 area 02 area 03 area 04 area 05 area 06 area 07 area 08 area 09 area 10 area 11 fig. 54 mgo vs. cao/al2o3 (wt%). data points outline a cluster showing decreasing mgo with decreasing cao/al2o3 with fractionation of both plagioclase and clinopyroxene (mayborn 2000). a number of analyses with mgo > 5% fall below trend line (1) and come from e–w-trending dykes (mayborn 2000; areas 05 and 06) and from e–w-trending dykes in the northern areas (area 08–11). (1): trend line 1. (2): trend line 2. 0.0 0.5 1.0 1.5 0.0 2.5 5.0 7.5 10.0 mgo c ao /a l 2o 3 (1) (2) area 01 area 02 area 03 area 04 area 05 area 06 area 07 area 08 area 09 area 10 area 11 https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 48 of 96 geusbulletin.org the immobile elements (right) is typical of other dyke swarms in continental settings and flood basalts (e.g. cadman et  al. 2001). this ‘crustal’ signature is distinct from that of more depleted normal mid-ocean ridge basalts (n-morb) and even transitional mid-ocean ridge basalts (t-morb) from other oceanic settings (e.g. rollinson 1993). however, the signature is quite similar to those of more enriched and transitional morbs (e-morb and t-morb, respectively; see fig. 56a) with one exception being the negative nb anomaly seen in all the dykes. nb depletion typically reflects subduction-related settings, but as that is unlikely for the kangâmiut dykes, the signature must stem from interaction with the continental crust. the consistency of the patterns for both mobile and immobile elements is also an indication that regionally, dykes have escaped significant element remobilisation as a result of metamorphism (except for rb in area 11, see chapter 7). in addition, the similarity to e-morbs strongly suggest an origin in a rift-like tectonic environment in a constructive plate-margin setting. in chapter 7, we address any compositional changes that may have accompanied the penetrative amphibolite and granulite facies metamorphism of dykes and their country rocks in the northern part of itillip ilua – ikertooq (areas 08–11) as well as the implications of the tectonic settings alluded to by the trace element patterns. the geochemical uniformity among the members of the kangâmiut dyke swarm is also evident in the rare earth element (ree) patterns (fig. 56b). of the 203 samples with ree data, we have selected 145 that had a complete set of data (from la to lu). additionally, we omitted nine samples, from which field evidence may not belong to the main kangâmiut swarm. the chondrite-normalised patterns exhibit very shallow to moderate slopes. for most samples, the steeper slopes are commonly associated with greater enrichment in the light ree (lree; fig. 56, left side) but only slight increase in the heavy ree (hree; fig. 56, right side). the magnitude of enrichment appears to be correlated with the location within the dyke where the sample was collected. for a few dykes, sampling traverses from margin to centre were available, and the samples collected near dyke margins consistently showed lower enrichments of lree than samples collected near dyke centres. this pattern is consistent with crystallisation of the dyke initiating at the dyke margins with no or minor enrichment, and slight to moderate progressive enrichment towards the centre, as incompatible lrees become more concentrated in the residual magma. as shown by mayborn et al. (2008), the dykes with leucocratic centres of andesitic composition typically displays such enrichment patterns. the nine samples omitted from this analysis had consistently steeper slopes that cross-cut the slopes of the remainder of samples. these samples are compositionally distinct from the other dykes, are consistently more mg-rich and likely represent separate intrusions. the overall gently sloping ree pattern of most dykes parallels that of classic e-morb (saccani et al. 2004; hémond et al. 2006; li et al. 2015). as shown in fig. 56b, the ree patterns for the kangâmiut dyke data set fig. 55 chemical variations for zr (ppm) and mg# from margins to centres of selected kangâmiut dykes. mg#: magnesium number, calculated as [mgo]/([mgo]+[feo(tot)]) × 100, where [ ] indicates molecular proportions. dyke orientation (black lines) and thickness (in m) are shown where available. data sources: h: hansen (1989). w: windley (1970). b: bridgwater (unpublished data). m: mayborn (2000). k: korstgård (1980). each block of data separated by vertical stippled lines represents a single dyke sampled from margin (left) to centre (right). some dykes show normal fractionation trends from margin to centre (decreasing mg# and increasing zr, for example the first dyke in area 02 and the rightmost dyke in area 04). other dykes show a more asymmetric distribution where most fractionated material is encountered away from the centre (for example the second dyke in area 04). remaining dykes show no signs of any margin to centre variations (for example the first dyke in area 05). for all dykes and samples, mg# and zr show opposite trends. 0 –50 –100 –150 –200 –250 –300 –350 –400 0 10 20 30 40 50 60 40 m 65 m 60 m 40 m 60 m 50 m 100 m 45 m 120 m 60 m 30 m zr (ppm)mg# h w h h h h b bb mh h h h h h k area 02 area 03 area 04 area 05 06 07 08 https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 49 of 96 geusbulletin.org fig. 56 trace and rare earth element (ree) characteristics of kangâmiut dykes. data sources in fig. 46 and supplementary file s1. a: multi-element plot for kangâmiut dykes from areas 01 to 11 normalised to pm (primitive mantle; values from mcdonough & sun 1995). for simplicity, only mean values of analyses from each area are displayed. light grey shading: range of all data (n = 239). shown are the mean values of enriched-, transitionaland normal-mid-ocean ridge basalt (e-morb, t-morb, n-morb, respectively, from hémond et al. 2006). a strong overlap with e-morb is noted. b: ree plot of kangâmiut dykes normalised to chondrite (normalisation values from mcdonough & sun 1995). rather than displaying all 136 samples, we show the density distribution (dark grey shading) for each normalised ree, thus clearly representing the wider range of values for light ree (lree) and narrow range for heavy ree (hree). red diamonds: mean of each normalised element. green diamonds: median. light grey shading: range of all data (n = 136). the mean and median values of lrees diverge according to the variable enrichment of these elements. the data showed no discernible variations between areas, so they are displayed together here. e-morb and t-morb are individual analyses from hémond et al. (2006) and waters et al. (2011), respectively. a ba rb th k nb la ce sr nd p zr sm eu ti tb y yb tm lu 0.1 1 10 100 sa m pl e/ pm area 01 area 02 area 03 area 04 area 05 area 06 area 07 area 08 area 09 area 10 area 11e-morb mean n-morb mean t-morb mean 1 10 100 la ce pr nd sm eu gd tb dy ho er tm yb lu sa m pl e/ c ho nd ri te b e-morb t-morb ree mean ree median https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 50 of 96 geusbulletin.org are indistinguishable from those of certain e-morb samples from the mid-atlantic ridge and show some overlap with t-morb (hémond et al. 2006). saccani et al. (2004) argued that such characteristics in basaltic dykes intruded into continental crust are consistent with a tectonic model of rifting involving magma genesis affected by mantle plume interaction with morb-type asthenospheric mantle. similar to that shown for the trace element distribution described here, the strong correlation of the kangâmiut dyke ree patterns with e-morb lends  support to a model in which the kangâmiut dykes represent the magmatic emplacement of rift  volcanics during continental extension or break-up. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 51 of 96 geusbulletin.org 5 metamorphic overprints on kangâmiut dykes from maniitsoq to itillip ilua the structural and metamorphic evolution of the kangâmiut dykes and their host rocks in the region between itillip ilua and ikertooq differs in many respects from the changes observed between maniitsoq and itillip ilua and will therefore be described separately (chapter 6). previous studies described the kangâmiut dykes in the southern foreland as largely pristine and unaffected by nagssugtoqidian orogenesis, based on the field appearance of the dykes. the widespread occurrence of garnets was assigned to thermal effects from the slowly cooling basic magma (e.g. jack 1978). ramberg (1949) and noe-nygaard & ramberg (1961b) only briefly describe the conversion of the kangâmiut dykes to schistose amphibolites along kangerlussuaq. dykes from the maniitsoq area, consisting of dolerite at the margins and garnet amphibolite in the centre of larger dykes, are described by windley (1970). in the inussuttusup tunua area (fig. 2), the metamorphic overprints are recognised as being more significant. jack (1978) described a suite of corona structures in the kangâmiut dykes in detail and concluded that these metamorphic features were a result of “autometamorphism”, as the dykes cooled slowly from their consolidation temperature (c. 900°c) to ambient country rock temperatures (500–700°c). we demonstrate that there is not a gradual increase in metamorphic grade recorded by the dykes from maniitsoq to itillip ilua. rather that most of the features listed here, from minor hydration of igneous assemblages to partial conversion of dykes to amphibolite facies amphibolites and garnet amphibolites can be observed throughout that area. only in the area north of itillip ilua (fig. 2) does nagssugtoqidian metamorphism and deformation result in the gradual to wholesale conversion of dykes to amphibolites, culminating in ikertooq where the metamorphic grade recorded by the dykes and their hosts reach granulite facies. the descriptions in the following sections are arranged in order of increasing signs of metamorphic overprints (i.e. any post-emplacement modifications of the original igneous mineralogy) as observed across the study area. the fully recrystallised dykes in amphibolite and granulite facies from saqqap kangerluarsua to maligiaq (fig. 2) are treated in chapter 6. 5.1 amphibole rims on ilmenite one of the first signs of hydration of the igneous mineral assemblage (clinopyroxene, plagioclase, ilmenite, ±amphibole, ±orthopyroxene, see also section 4.2) is the formation of thin, blue-green amphibole rims on ilmenite. figure 56 shows various stages of the development of amphibole rims. these rims typically occur where ilmenite is in contact with plagioclase, whereas ilmenites in contact with clinopyroxene may not show amphibole rims. compositionally, these amphiboles are “ferro-pargasites” (following the nomenclature of hawthorne et al. 2012; see also fig. 48a). adjacent plagioclase is clearly involved in the amphibole-forming reactions, as anorthite content in plagioclase (an) is lowered drastically along the amphibole rims (an20–25 vs. an60 in cores of igneous plagioclase). as seen in fig. 57, these amphibole rims occur both in fine-grained samples from marginal chills and from coarser-grained parts of dykes with well-developed ophitic textures. it is notable, that in samples with only amphibole rims on ilmenites, the clinopyroxenes appear fresh, and have not been replaced, even partially, by amphiboles. 5.2 amphibole rims on clinopyroxene further hydration is recorded by the formation of a more extensive network of blue-green to green amphibole rims on igneous clinopyroxenes (fig. 58). these rims are never seen in microstructural settings where adjacent ilmenites have no rims and are thus interpreted to represent a further step in the metamorphic development. as in the case of amphibole rims on ilmenites, the participation of plagioclase in this amphibole-forming reaction is shown by the lower an contents in adjacent plagioclase, and the disturbance of the igneous zoning patterns as seen in the microscope. the continued growth of amphibole in the rims appear to largely take place at the expense of plagioclase (fig. 58). these amphibole rims also seem to develop on clinopyroxenes that have undergone partial or full replacement by amphiboles (fig. 58b, c), suggesting that the alteration of clinopyroxenes was initiated prior to rim development. igneous orthopyroxenes are more resistant to alteration than coexisting clinopyroxenes, and also appears to have a less well-developed amphibole rims (fig. 58b). https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 52 of 96 geusbulletin.org 5.3 partial to complete replacement of clinopyroxene by amphibole along with the growth of amphibole along the margins, the internal parts of clinopyroxenes may be partially to completely replaced by amphiboles, typically with a greenish colour (various examples shown in fig. 59). replacement occurs initially along cracks, fractures and internal grain boundaries in clinopyroxene aggregates. ilmenite grains may be seen in or near the boundary fig. 57 early metamorphic overprints of amphibole rims on ilmenite. a: chilled margin of 30 m wide kangâmiut dyke. primary amphibole and clinopyroxene overgrowing a fine-grained matrix of plagioclase, clinopyroxene and ilmenite. example of incipient formation of thin amphibole rims on ilmenite shown in circle. 1 nicol. sample 430267. south shore, mouth of kangerlussuatsiaq. b: margin of 40 m wide dyke with pronounced podding geometry. brown-green amphibole overgrowing ilmenite grains, mainly at the expense of adjacent plagioclase. neighbouring clinopyroxenes show no amphibole rims. 1 nicol. sample 414304. north shore, mouth of itillip ilua. c: fine-grained ne-trending dyke sampled a few centimetres from contact to older, coarse-grained e–w-trending dyke (outcrop shown in fig. 10c). blue-green amphibole rims on ilmenites. 1 nicol. sample 430274. south shore, mouth of kangerlussuatsiaq. amp: amphibole. cpx: clinopyroxene. ilm: ilmenite. pl: plagioclase. ilm a b c 0.5 mm 0.5 mm 0.5 mm 430267 amp amp cpx cpx ilm 414304 430274 cpx pl pl pl ilm ilm a b c 0.5 mm 0.5 mm 0.5 mm cpx cpx opx pl pl pl cpx cpx amp amp amp amp amp amp cpx ilm 158027 158035 158027 fig. 58 early metamorphic overprints: amphibole rims on ilmenite and clinopyroxene. a: green amphibole rims on ilmenite (top right) and clinopyroxene. locally, the plagioclase rims adjacent to amphibole rims are brighter and free from inclusions. centre of nne-trending dyke. 1 nicol. sample 158027. southern inussuttusup tunua. b: thin, green amphibole rims on ilmenite (centre), clinopyroxene and orthopyroxene. centre of nne-trending dyke. 1 nicol. sample 158027. southern inussuttusup tunua. c: thin, blue-green amphibole rims on ilmenite and clinopyroxene. locally, amphiboles also grow on internal boundaries in clinopyroxene aggregates. amphiboles may or may not grow along ilmenite–clinopyroxene interfaces, possibly reflecting lack of access to necessary reactants (i.e. fluid, plagioclase). 1 nicol. sample 158035, 2 km north of kangaamiut. amp: amphibole. cpx: clinopyroxene. ilm: ilmenite. opx: orthopyroxene. pl: plagioclase. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 53 of 96 geusbulletin.org between newly formed amphibole and clinopyroxene, indicating that ilmenite participated in the amphibole-forming reaction. in the final stages, the entire grain is replaced with a quartz inclusion-filled, poikilitic, green to dark green amphibole (fig. 59d). the original amphibole rim, which initially surrounded the igneous clinopyroxene, is in many cases still visible after complete replacement of the inner parts (fig. 59d). where the initial amphibole rim is typically composed of thin, radiating crystals perpendicular to the clinopyroxene grain boundary (see fig. 58), the later rim is often less distinct and recrystallised to a polygonal layer (fig. 59c). as mentioned above, the relative timing and progress of clinopyroxene conversion to amphibole varies significantly between samples, and even within a thin section. in some samples, well-developed garnet rims occur on partially converted clinopyroxenes, whereas in others, the early thin amphibole rims surround completely amphibolitised clinopyroxene grains or aggregates. we suspect that subtle variations of fluid accessibility and pressure–temperature (p–t) conditions are responsible for these varied developments of whole vs. partial amphibolitisation. 5.4 epidote growth in plagioclase epidote growth in igneous plagioclase is mainly observed in the southern part of the study area, however, as the epidotes vary substantially in both size and crystal form, it is possible that very small grains were occasionally overlooked under the microscope and are thus more widespread. well-developed, large (up to 0.3 mm) lath-shaped crystals develop in the central, more ca-rich parts of the igneous plagioclase (fig. 60a, b). where subsequent recrystallisation of the plagioclase has occurred, the clusters of epidotes mimic the original shapes of the igneous plagioclases (fig. 60b). the brownish dusting that is commonly observed in igneous plagioclases, presumably disseminated fe oxides (see figs 57, 58), a b c 0.5 mm 0.5 mm 0.5 mm cpx cpx amp amp pl amp + qtz amp-rim grt cpx amp -1 amp-2 cpx cpx d 0.5 mm 430275 87777414445 414429 amp-rim fig. 59 early metamorphic overprints: partial to complete replacement of clinopyroxene by amphibole. a: centre of 1 m wide pod in kangâmiut dyke. green amphibole grows on all clinopyroxene–plagioclase and clinopyroxene–clinopyroxene contacts. 1 nicol. sample 414429. equalugaarsuit, north shore itillip ilua. b: coarse-grained part of 50 m wide n–s-trending dyke. clinopyroxene aggregate, partially replaced by pale-green amphiboles along rims and along internal boundaries. 1 nicol. sample 430275. south shore, mouth of kangerlussuatsiaq. c: 70 m wide, foliated, recrystallised and zoned ne-trending dyke. clinopyroxene with rims of green amphibole (amp-1) and advanced internal replacement by lighter-green amphiboles (amp-2). 1 nicol. sample 414445. south shore, mouth of itillip ilua. d: 48 m wide kangâmiut dyke showing pronounced compositional zoning. sample from the central part of dyke. advanced replacement of clinopyroxene by recrystallised, green amphibole and quartz in a poikilitic aggregate. euhedral garnets grow in parts of the amphibole rim. 1 nicol. sample 87777. coast just north-east of sermersuut. amp: amphibole. amp-rim: amphibole rim. cpx: clinopyroxene. grt: garnet. qtz: quartz. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 54 of 96 geusbulletin.org tends to disappear in the vicinity of epidotes and may thus take part in the hydration reaction. compositionally, the epidotes show only minor variation and are either ‘epidotes’ or ‘clinozoisites’, depending on whether fe3+ or al3+, respectively, is the dominant cation in the m3 site (following the classification of armbruster et al. 2006). 5.5 garnet growth on ilmenite and in amphibole rims on ilmenite coincident with the early stages of amphibole rims forming on clinopyroxenes, garnets begin to form along ilmenite rims (various stages of reaction progress shown in fig. 61). in most cases, garnet growth is clearly replacing a pre-existing amphibole rim (fig. 61), whereas in other cases, garnet develops directly at the ilmenite:plagioclase interface, with no remnants of amphibole. considering that ilmenites without amphibole rims are rare among the least metamorphosed dykes, it is thought that garnet rims on ilmenite without amphibole reflect complete consumption of an earlier amphibole. early garnets occur as fine-grained, inclusion-filled aggregates, and as garnet growth progresses in either setting, grains are increasingly inclusion-free and euhedral, especially in their outer rims against adjacent plagioclase. figure 61d shows an example of a complex garnet rim along half the perimeter of an ilmenite. in this example the inner part of the rim is riddled with elongate grains of biotite and amphibole, all perpendicular to the ilmenite grain boundary, while the outer part consists of clear, inclusion-free euhedral garnet grains. along the remainder of the ilmenite grain boundary, the amphibole rim has yet to be replaced. compositionally, the garnets show minimal variation across the rim, while the varying mineralogy inside the rims suggests that a continuum of mineral reactions is responsible for the rim assemblages. at higher metamorphic grades, recorded in the area around itillip ilua, the garnet rims on ilmenites become even more complex. figure 62 shows an example of complex garnet rims on some parts of the ilmenite grains, while other parts of the ilmenites and the adjacent clinopyroxene only display amphibole rims. 5.6 garnet growth in amphibole rims on clinopyroxene further progress of metamorphic overprinting is recorded by the growth of garnets in the amphibole rim surrounding clinopyroxene (fig. 63, increased reaction progress from a to c). most observations suggest that the garnet initially grows at the clinopyroxene-amphibole interface, mainly at the expense of amphibole. however, other geometries, for example garnet replacing the amphibole rim from the amphibole-plagioclase grain boundary and inwards, are also observed (fig. 63b). where amphiboles have been completely consumed, garnet grains form ‘atolls’ consisting of a b c 0.5 mm 0.5 mm 0.5 mm 430294 430294 430285 pl + ep pl + ep pl + ep pl + ep ilm+amp ttn-rim amp amp grt grt grt amp fig. 60 early metamorphic overprints of epidote growth in plagioclase. a: coarse-grained, central mafic part of 40 m wide composite kangâmiut dyke. randomly orientated epidote laths in plagioclase. the density of epidote outlines compositional zoning patterns in original, igneous plagioclase. epidote-rich plagioclases separated by inclusion-free plagioclase. 1 nicol. ungusivik island, 8 km south of mouth of kangerlussuaq. b: same sample as panel a. here, garnet has entered the assemblage. epidote clearly outlines compositional zoning pattern in plagioclase. top left, skeletal ilmenite in amphibole suggests advanced replacement of ilmenite. thin rims of titanite can be seen on the skeletal ilmenite. c: garnet-bearing centre of 30 m wide ne-trending dyke. euhedral garnet growing between plagioclase and amphibole, largely at the expense of plagioclase. epidote needles in plagioclase show a wide range in size (0.01–0.1 mm). 1 nicol. south shore, mouth of kangerlussuatsiaq. amp: amphibole. ep: epidote laths. grt: garnet. ilm: ilmenite. pl: plagioclase. ttn-rim: titanite rim. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 55 of 96 geusbulletin.org euhedral crystals, typically protruding into, and sometimes almost replacing the plagioclase (fig. 63c). as noted earlier, the participation of plagioclase in these reactions is shown by the significant drop in an content adjacent to the garnets. 5.7 titanite rims on ilmenite in a few instances, narrow titanite rims have been observed on ilmenite (fig. 60b). these ilmenites may display a pronounced skeletal habit (fig. 64a), presumably reflecting the involvement of ilmenite in the amphibole and garnet forming mineral reactions as well as the formation of titanite. skeletal ilmenites are also observed in situations where titanite rims are not developed and may similarly reflect the involvement of ilmenite in other metamorphic reactions. the examples in figs 64a, b show skeletal ilmenites, with titanite overgrowths that in turn are overgrown by garnets, thus suggesting that the ilmenite–titanite transition predated garnet growth. a b c d 0.5 mm 0.5 mm 0.5 mm 0.5 mm grt-rim amp-rim amp-rim amp-rim amp-rim amp-rim ilm ilm cpx cpx cpx cpx pl pl grt-rim cpx>>amp cpx>>amp grt-rim grt-rim 158035 158035 414399 158035 fig. 61 early metamorphic overprints of garnet growth on ilmenite. a: delicate garnet rims along ilmenite–plagioclase interfaces and pale green amphibole rims between clinopyroxene and plagioclase. 1 nicol. sample 158035. 2 km north of kangaamiut. b: same as panel a. thin garnet rims between ilmenite and plagioclase. amphibole rims on most clinopyroxene–plagioclase contacts. around the perimeter of a plagioclase grain, garnet is encountered at ilmenite contacts, and amphibole at clinopyroxene contacts. generally, no garnet or amphibole occur along ilmenite–clinopyroxene contacts. 1 nicol. sample 158035. 2 km north of kangaamiut. c: similar to panels a and b, except that in this domain, clinopyroxenes are being replaced by amphiboles, both internally and along aggregate boundaries (cpx>>amp). thin blue-green amphibole rims are developed between ilmenite and clinopyroxene. 1 nicol. sample 158035, 2 km north of kangaamiut. d: margin of 10 m wide e–w-trending dyke in itillip ilua. ilmenite has a green amphibole rim, which is being replaced locally by a garnet rim consisting of an inner, inclusion-filled rim, and an outer, euhedral rim protruding into neighbouring plagioclase. 1 nicol. sample 414399. eqalugaarsuit, north shore, itillip ilua. amp: amphibole. cpx: clinopyroxene. grt: garnet. ilm: ilmenite. pl: plagioclase. fig. 62 a 70 m wide, foliated, recrystallised and zoned ne-trending dyke. two ilmenite grains with complex garnet rims. ilmenite to the right has an amphibole rim along its right side and a garnet rim along the rest of its perimeter. inner garnet rim contains very fine-grained, radiating minerals (biotite, ilmenite), whereas the outer rim is inclusion-free and shows euhedral crystal faces protruding into plagioclase. the ilmenite to the left shows incipient skeletal texture in the upper right part of the grain and has a double rim of garnet towards the bottom and left. inner rim and inner part of the outer rim contain thin, radiating inclusions, while the outer rim is inclusion-free with euhedral grains. adjacent clinopyroxene has a recrystallised amphibole rim, but no garnets. 1 nicol. sample 414445. south shore, mouth of itillip ilua. amp: amphibole. cpx: clinopyroxene. grt: garnet. ilm: ilmenite. 0.5 mm cpx ilm ilm grt-rims amp-rim amp-rim amp-rim 414445 https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 56 of 96 geusbulletin.org 5.8 disseminated garnet in dyke margins in the fine-grained margins of some dykes, garnet-rich zones sometimes occur along what appears to be cracks or fractures, typically at a high angle to the dyke contact. the zones can be millimetre to centimetre in scale and can sometimes show curvature towards the margin reflecting an apparent sense of shear. figure 65a shows an example of garnet-filled millimetre-scale veins at high angles to the contact. towards the centre of the dyke, the bands are wider (millimetre–centimetre scale) and more diffuse (fig. 65b). the central part of many dykes is coarser-grained garnet-amphibolite (fig. 65c). in addition to garnet-filled veins, some dykes also show numerous 5–10 mm wide garnet clusters throughout their contact zones (see the lighter ‘spots’ in the darker bands in fig. 65a). in thin section, these garnet-filled veins are relatively fine-grained (fig. 66a) and composed of garnet and ilmenite and a very fine-grained matrix. the veins appear to truncate the primary subophitic texture in the dyke margins without much interaction with the igneous mineralogy (fig. 66a), at times surrounding ophitic igneous clinopyroxene aggregates entirely. where coarser-grained dykes are invaded by the garnet-filled veins, the resultant texture is more diffuse, and the veins seem to advance along boundaries of the larger plagioclase and clinopyroxene grains (fig. 66b). the timing of the formation of these garnet-rich veins, layers and clusters is most likely coincident with the overall metamorphic overprint seen in the dykes and it is thought that the garnet-filled veins utilised pre-existing structures and weaknesses as conduits when they propagated into the dykes. 5.9 distribution of metamorphic overprints from maniitsoq to itillip ilua the metamorphic features described above are schematically summarised in fig. 67, which illustrates the mineralogical changes that occur at the igneous ilmenite–plagioclase and clinopyroxene–plagioclase boundaries, respectively. the progress of metamorphic overprints increases from left to right (more information in figure caption). in terms of reaction progress, it is notable, that reactions at the ilmenite-plagioclase boundaries appear to proceed before those at the clinopyroxene-plagioclase interfaces. figure 68 shows the regional distribution of the metamorphic features described here and in fig. 67. this compilation also includes a few observations from hansen (1989) and from mayborn (2000). as mentioned previously, the area from maniitsoq to itillip ilua does not represent a simple, gradual progress in the metamorphic overprints, as schematically shown in fig. 67. rather it demonstrates that most of the features can a b c 0.5 mm 0.5 mm 0.5 mm cpx cpx cpx ilm ilm pl pl pl pl grt grt grt amp grt cpx grt-rim 414409 414445 414411 fig. 63 early metamorphic overprints of garnet growth in amphibole rims on clinopyroxene. a: centre of 1.5 m wide dyke. advanced growth of garnet on ilmenite. amphibole rims on all clinopyroxene, and locally garnet, partly or entirely replaces amphibole. outer part of most garnet rims shows inclusion-free, euhedral growths into plagioclase. 1 nicol. sample 414409. north shore, central part of itillip ilua. b: 70 m wide, foliated, recrystallised and zoned ne-trending dyke. wide garnet rims on ilmenite and garnet growth in amphibole rims surrounding clinopyroxenes. 1 nicol. sample 414445. south shore, mouth of itillip ilua. c: garnet-bearing, leucocratic zone in coarse-grained dyke. advanced growth of garnet on ilmenite and clinopyroxenes. outermost rims are inclusion-free and euhedral. sample 414411. north shore, central part of itillip ilua. amp: amphibole. cpx: clinopyroxene. grt: garnet. ilm: ilmenite. pl: plagioclase. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 57 of 96 geusbulletin.org be found throughout the area, clearly suggesting that whatever drove these metamorphic processes (pressure, temperature, availability of fluids) was heterogeneously distributed throughout this southern foreland region. the metamorphic drivers, their causes and the conditions under which they operated are further discussed in chapter 9. a quantitative look at the variations of both igneous and metamorphic minerals from maniitsoq to itillip ilua is illustrated in fig. 49, which represents a compilation of all available modal analyses. despite variability in terms of the detail and number of points analysed, and in accuracy when discriminating between igneous and metamorphic minerals, it is noteworthy how relatively little variation there is in the contents of both felsic (plagioclase + quartz) and mafic (pyroxenes + amphibole) minerals from area 02 to area 07. this underscores an earlier assertion based on bulk-rock a b 0.5 mm 0.5 mm414432 414432 ilm grt amp grt ttn-rims ttn fig. 64 early metamorphic overprints of titanite rims on ilmenite. a: garnet-bearing, leucocratic material from the centre of a large dyke. skeletal ilmenite partially replaced by titanite along parallel planes. ilmenite–titanite aggregate (red dashed line) completely overgrown by garnet. 1 nicol. sample 414432. north shore, mouth of itillip ilua. b: same sample as panel a, but with much more irregular geometry of titanite replacement and overgrowth of ilmenite. 1 nicol. sample 414432. north shore, mouth of itillip ilua. amp: amphibole. grt: garnet. ilm: ilmenite. ttn: titanite. a b c a b c fig. 65 early metamorphic overprints of disseminated garnet in margins of dykes. a: sharp contact between kangâmiut dyke (upper part of image) and folded leucocratic host gneiss (lower part). broadly parallel, non-continuous 1–2 cm wide, lighter, garnet-rich bands can be seen at a high angle to the dyke contact. within darker bands, millimetre-scale, garnet-rich aggregates are widely distributed and appear as lighter grey dots. eqalugaarsuit, north shore itillip ilua. b: fine-grained dyke with primary compositional layering parallel to margin and accentuated by alternating garnet-rich (lighter brown bands) and amphibole-richer (darker colour) zones. eqalugaarsuit, north shore itillip ilua. c: detail of homogeneous, isotropic and coarse-grained part of dyke, made up of plagioclase (light coloured), amphibole (dark grey) and garnet (dark red specs). eqalugaarsuit, north shore itillip ilua. scale is 10 cm in all the figures. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 58 of 96 geusbulletin.org ilmenite plagioclase plagioclase clinopyroxene garnet garnet plagioclase 1 plagioclase 1 amphibole 1 amphibole 2 amphibole fig. 67 schematic to illustrate the metamorphic reactions that take place at ilmenite–plagioclase (top) and clinopyroxene–plagioclase (bottom) interfaces, with increasing degree of reaction progress from left to right. note that both amphibole and garnet formation occur earlier along ilmenite than along clinopyroxene margins. plagioclase: original igneous plagioclase with anorthite contents up to an60. plagioclase-1: metamorphic plagioclase, initially participates in the amphibole-forming reactions and later in the garnet-forming process. plagioclase-1 has much lower anorthite content (an20-30). amphibole and amphibole-1: first grow along ilmenite and clinopyroxene margins, respectively. amphibole-2 is the result of replacement of the internal parts of clinopyroxene grains by amphiboles, typically along fractures, cleavage planes or in patches. not shown is the epidote formation within plagioclases and titanite growth at the expense of ilmenite. symbols above and below the diagram corresponds to the mineral assemblage symbols in fig. 68. a b 0.5 mm 0.5 mm414433 414443 pl cpx amp amp cpx grt+ilm cpx cpx pl grt+ilmamp -rim fig. 66 early metamorphic overprints of disseminated garnet in margins of dykes. a: fine-grained dyke margin with garnet-filled veins and bands. partially altered clinopyroxene with recrystallised rims of brown-green amphibole. irregular, discordant band of fine-grained aggregate of garnet and ilmenite cuts across the thin section (top right to lower left) 1 nicol. sample 414433. north shore, mouth of itillip ilua. b: margin of nne-trending and 40 cm wide dyke. igneous minerals and texture are largely preserved. a few clinopyroxenes have thin green amphibole rims. in diffuse, fine-grained aggregate (1–2 mm across), garnets overgrow all ilmenites and line most other grain boundaries. 1 nicol. sample 414443. south shore, mouth of itillip ilua. amp: amphibole. cpx: clinopyroxene. grt: garnet. ilm: ilmenite. pl: plagioclase. chemistry, that the igneous precursors to these variably metamorphosed dykes were compositionally homogeneous on a regional scale. the modal variations shown in fig. 49 are interpreted to reflect the very heterogeneous progress of reactions leading to the growth of amphibole rims on igneous minerals to the complete replacement of igneous phases. in areas 09–11, all minerals are metamorphic. the volume of garnet increases from south to north and reaches a maximum of c. 10 vol.% in areas 06 and 07, but then drastically decreases farther north, only to sporadically pick back up again in area 11. in areas 08–11, the structural and metamorphic evolution of kangâmiut dykes and their host rocks differs from the changes observed between maniitsoq and itillip ilua and are treated separately next. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 59 of 96 geusbulletin.org 52° 52° 67° 67° 66° 54° 66° 54° sisimiut kangaamiut maniitsoq ilm w/amp-rim ilm w/amp-grt rim ilm w/grt rim cpx w/amp-rim cpx w/amp-grt rim pl w/ep-incl 50 km fig. 68 simplified map of the mineral assemblages described in section 3.1 showing the distribution of observed metamorphic overprints at individual sample locations. see fig. 1 for explanation of fault and dyke symbols and other place names. mineral abbreviations follow whitney & evans (2010). ilm w/amp-rim: ilmenite with amphibole rims. ilm w/amp-grt rim: ilmenite with rims of amphibole and garnet. ilm w/grt rim: ilmenite with garnet rims. cpx w/amp-rim: clinopyroxene with amphibole rims. cpx w/amp-grt rim: clinopyroxene with rims of amphibole and garnet. pl w/ep-incl: plagioclase with epidote inclusions. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 60 of 96 geusbulletin.org 6 deformation and metamorphism of dykes and country rocks from saqqap kangerluarsua to maligiaq 6.1 microstructural changes the conversion of the dolerite dykes into metadykes or amphibolites can be observed south of saqqap kangerluarsua and especially in qaqqatoqaq and along itillip ilua (fig. 35). here, every stage from unaltered dolerite dyke to metadyke is exposed. the conversion took place under varying degrees of deformation and although the final product is essentially the same, apart from variations in grain size and intensity of linear and planar fabric, there are some important differences in the paths taken by the dykes during their conversion into amphibolites. we tend to use metadyke and amphibolite synonymously, since metamorphosed basaltic dykes per definition are amphibolites. however, some amphibolites may have a different precursor (e.g. leake 1964). in this section, the dyke-to-metadyke conversion will be described in terms of accompanying changes in microstructure, including changes in shape and arrangement of crystals (vernon 2004; vernon & clarke 2008), from a non-metamorphic microstructure to a granoblastic metamorphic microstructure. two cases are considered: (1) microstructural changes in dykes metamorphosed during weak deformation, and (2) microstructural changes in dykes during strong deformation. deformation here refers to deformation as observed in thin sections and not necessarily to deformation as seen from the geometry of the dyke body. 6.1.1 dykes metamorphosed under weak deformation in weakly deformed dykes, the changes in shape and arrangement of grains are exclusively associated with a change in the mineralogy of the dykes. the metamorphic overprints described in chapter 5, mainly growth of amphibole and garnet, has little effect on the microstructure as the original igneous texture is still recognisable. further growth of amphibole and garnet and recrystallisation of plagioclase, however, eventually alters the microstructure from a sub-ophitic grain-contact arrangement to a granoblastic (polygonal) grain-contact arrangement. mineralogically speaking, the dykes are now amphibolites since they are composed of metamorphic minerals, even if the original sub-ophitic distribution of dark and light mineral phases may still be discernible. in the following sections, a few pertinent steps in this metamorphic transition are described and illustrated. 6.1.1.1 igneous microstructure the microstructures of undeformed and unmetamorphosed dykes range from fine-grained porphyritic to sub-ophitic or coarse-grained, almost granular. the porphyritic microstructure is usually found in the margins of larger dykes. the phenocrysts are typically an-rich and up to 1 mm long plagioclases. less frequently, smaller clinopyroxene phenocrysts with irregular outlines occur. the matrix is fine-grained (< 0.05 mm) and consists of clinopyroxene, plagioclase, ilmenite, and primary amphibole. examples of porphyritic microstructures are shown in fig. 69. the most common microstructure is a sub-ophitic microstructure where plagioclase crystals penetrate but are not enclosed by pyroxenes. in coarse-grained varieties, the sizes of the plagioclase crystals vary between 0.2 and 2 mm (fig. 70), whereas in fine-grained dolerites this range is 0.1–0.5 mm (fig. 71). 6.1.1.2 incipient granoblastesis, no deformation as described in chapter 5, most kangamiut dykes show signs of a metamorphic overprint, where amphiboles and garnets typically form at the expense of ilmenites and clinopyroxenes. initially, the amphiboles are small, granoblastic crystals without any preferred shape orientation or pleochroism. plagioclases becomes clouded (e.g. figs 57b, 58a–c, 69, 70) and small, clear, granoblastic plagioclase crystals may start to develop around plagioclase laths and along plagioclase-plagioclase grain contacts. throughout these mineralogical adjustments, the igneous sub-ophitic microstructure remains well preserved. 6.1.1.3 intermediate granoblastesis, no deformation at an intermediate stage of metamorphism, all clinopyroxene crystals are surrounded by rims of small, granoblastic amphibole crystals (fig. 72c). this gives the rock a darker appearance in hand specimen, while in thin section, sub-ophitic microstructures become less distinct (fig. 72a, b). at this stage, plagioclase crystals are surrounded by small new plagioclase crystals, ranging in size from 0.05 mm to 0.2 mm (fig. 72a, b). this stage is also characterised by widespread development of garnet, which has now grown to a size of c. 0.2 mm. while approximately half of the crystals in these rocks are granoblastic, the distribution of light and dark minerals is still reminiscent of the original sub-ophitic microstructure (fig. 72b). https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 61 of 96 geusbulletin.org fig. 69 examples of porphyritic microstructure in fine-grained dyke. a, b: porphyritic microstructure at dyke margin. c, d: phenocrysts of plagioclase in a matrix of clinopyroxene, plagioclase, primary amphibole and ilmenite. a, c: 1 nicol. b, d: x nicols. sample 1143. outer saqqap kangerluarsua. a b c d a b c d fig. 70 examples of sub-ophitic microstructure in coarse-grained dyke. a, b: coarse-grained dyke with sub-ophitic microstructure. c, d: laths of plagioclase with interstitial clinopyroxene. a, c: 1 nicol. b, d: x nicols. sample 1332. south shore of central itillip ilua. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 62 of 96 geusbulletin.org a b c d fig. 71 examples of sub-ophitic microstructure in fine-grained dyke. a, b: fine-grained dyke with sub-ophitic microstructure. c, d: laths of plagioclase with interstitial clinopyroxene. a, c: 1 nicol. b, d: x nicols. sample 1328. south shore of central itillip ilua. a b c d fig. 72 preserved sub-ophitic microstructure in dyke that is partly converted into a metadyke. a: 1 nicol. b: x nicols. c: 1 nicol, igneous clinopyroxene rimmed by amphibole. d: x nicols, small granular metamorphic plagioclase grains at igneous plagioclase interfaces. sample 1344. south shore of central itillip ilua. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 63 of 96 geusbulletin.org 6.1.1.4 advanced granoblastesis, no deformation all igneous clinopyroxene has disappeared at this advanced stage of metamorphism (fig. 73a, b). the amphiboles along the rims of former clinopyroxenes have now grown to larger granoblastic crystals, while their central part consists of aggregates of amphiboles and quartz (fig. 73c, d). most of the igneous plagioclase laths have also been replaced by smaller polygonal crystals, which have started to coalesce into larger crystals. in fine-grained dykes, mafic phases now consist of granoblastic aggregates of amphibole, garnet and ilmenite (fig. 74). a few igneous phenocrysts of plagioclase or clinopyroxene or both may survive this transformation but are strongly corroded. otherwise, these igneous laths are also converted into polygonal grains. sample 1286, in fig. 75, shows the final product of granoblastesis within a dyke with only a subtle indication of a planar fabric. although completely recrystallised, it clearly preserves the original sub-ophitic distribution of felsic and mafic phases. this is not as apparent in sample 1288 (fig. 75). 6.1.2 dykes metamorphosed under strong deformation the conversion of dykes into metadykes involves simultaneous deformation, as evidenced by the development of a strong directional fabric in the rocks, partially defined by metamorphic minerals. several steps, involving various degrees of deformation and recrystallisation could be outlined, but only two cases are described and illustrated below: (1) deformation involving original igneous minerals and (2) deformation accompanied by complete recrystallisation. 6.1.2.1 incipient granoblastesis, strong deformation a directional fabric is produced through the development of elongate aggregates of plagioclase and mafic minerals. most plagioclases break down to smaller equidimensional grains. small garnets and amphiboles develop together with ilmenite and form elongate trails that define the directional fabric (fig. 76). in contrast to dykes that have been metamorphosed without much deformation, the igneous clinopyroxenes within strongly deformed dykes often survive as clasts and may even recrystallise into smaller, polygonal aggregates that surround igneous relicts (fig. 77), a feature that is not observed within undeformed dykes. plagioclase – normally the best-preserved igneous relict in undeformed dykes – appears to break a b c d fig. 73 a dyke in which most clinopyroxene crystals have been replaced by amphibole. a: 1 nicol. b: x nicols. original igneous distribution of plagioclase still preserved. note sieve hornblende (aggregate of amphibole and quartz replacing clinopyroxene). c: 1 nicol. d: x nicols. sample 1319. north shore of central itillip ilua. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 64 of 96 geusbulletin.org down much more readily than clinopyroxene, when the rock is subjected to deformation. two examples of strongly deformed dykes are shown in fig. 78. these thin sections are cut both parallel to the strongest directional fabric and at right angles to this. it appears that the linear fabric is much more strongly developed than the planar fabric. it is also worth noting that, even in the most strongly deformed rocks in figs 77 and 78, the original igneous sub-ophitic microstructure is still discernible. 6.1.2.2 complete granoblastesis, strong deformation a combination of strong deformation and the disappearance of all igneous minerals effectively erases the original igneous microstructure. a transitional case is illustrated in fig. 79, where all minerals are metamorphic, but the distribution of plagioclase and amphibole is suggestive of an igneous origin. this is particularly evident in the thin section that is orientated perpendicular to the rock’s linear fabric (fig. 79a, b). with further growth of metamorphic minerals, the relict igneous mineral distribution is eventually lost. figure 80 shows this final development of a granoblastic, metamorphic fabric. the deformed nature of these rocks is indicated by a strong crystallographically preferred orientation of amphiboles as shown by their pleochroism (fig. 80c). a b c d ign cpx ign cpx fig. 74 a fine-grained dyke where most igneous minerals have been replaced by the metamorphic minerals, amphibole, garnet and plagioclase. some igneous clinopyroxene (ign cpx) grains have survived. a, c: 1 nicol. b, d: x nicols. sample 1316, north shore of central itillip ilua. a b fig. 75 a completely recrystallised dykes with a faint planar fabric. a: sample 1286. partly igneous distribution of mineral phases. b: sample 1288. igneous distribution of mineral phases can hardly be distinguished. both 1 nicol. qaqqatoqaq. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 65 of 96 geusbulletin.org 6.1.3 conclusions regarding granoblastesis microstructurally, the transformation from sub-ophitic dykes into granoblastic amphibolites involves (1) recrystallisation of existing minerals, (2) growth of new mineral phases and (3) the simultaneous migration of material to produce a rock where mineral phases are more or less evenly distributed or where these are concentrated in bands of light and dark minerals. the formation of a banded microstructure in the metadykes clearly involves strong deformation (fig. 78). igneous laths of plagioclase are transformed into elongate aggregates of smaller granoblastic grains through a b c d fig. 76 a strongly deformed dyke where lineation is much stronger than foliation. a (1 nicol), c: sample 1350a, cut parallel to the lineation. b (1 nicol), d: sample 1350b, cut perpendicular to the lineation. in c (x nicols), trails of fine-grained garnet (black) and recrystallised plagioclase outline the fabric. larger brownish grains survive igneous clinopyroxenes. d (1 nicol), cut across the linear structure, shows reddish garnet, ilmenite, a few green amphiboles and some recrystallised clinopyroxenes. south shore of central itillip ilua. a b c d e f m m z x x fig. 77 a strongly deformed dyke. sections cut perpendicular to the foliation. a: 1 nicol. b: x nicols. c–f: 1 nicol. ‘x’ in panels c, e: relic igneous clinopyroxene. ‘z’ in panel f: recrystallised clinopyroxene. sample 1345. south shore of central itillip ilua. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 66 of 96 geusbulletin.org a b c d fig. 78 a strongly deformed dyke. a: sample 1300, 1 nicol. b: sample 1300, x nicols. in b, garnets stand out as black. c: sample 1321a, 1 nicol, cut perpendicular to apparent foliation. d: 1321b, 1 nicol, cut perpendicular to lineation and showing the linear nature of the fabric. both samples north shore of central itillip ilua. a b c d e f m m z x x fig. 77 (continued) a strongly deformed dyke. sections cut perpendicular to the foliation. a: 1 nicol. b: x nicols. c–f: 1 nicol. ‘x’ in panels c, e: relic igneous clinopyroxene. ‘z’ in panel f: recrystallised clinopyroxene. sample 1345. south shore of central itillip ilua. a b c dfig. 79 a completely recrystallised dyke. a, c: sample 1141a, 1 nicol, perpendicular to lineation. b, d: sample 1141b, x nicols, parallel to lineation. outer saqqap kangerluarsua. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 67 of 96 geusbulletin.org dynamic recrystallisation. in response to the initial deformation, plagioclase laths develop new grains either marginally or along plagioclase–plagioclase grain contacts (fig. 72). as deformation proceeds, these aggregates are stretched into extremely long trails that conform to the prevailing strain field. this observation bears a striking resemblance to structures found in mylonites. 6.2 metamorphism of dykes and country rocks how undeformed dykes from maniitsoq to itillip ilua have been affected by metamorphic overprinting is covered in detail in chapter 5. in parts of the itillip ilua region, and especially north of qaqqatoqaq, these overprinting processes have run to completion and have resulted in dykes being converted into metadykes or essentially amphibolites. textural aspects of the transformation were covered in section 6.1. in this section, we describe the metamorphic effects of the nagssuqtoqidian deformation on dykes and country rocks north of itillip ilua. along the north shore of saqqap kangerluarsua and in qeqertalik (fig. 81), the effects of nagssugtoqidian deformation on the country gneisses are relatively slight and are mainly expressed as the folding of earlier fabrics, around e–w a b c d fig. 79 (continued) a completely recrystallised dyke. a, c: sample 1141a, 1 nicol, perpendicular to lineation. b, d: sample 1141b, x nicols, parallel to lineation. outer saqqap kangerluarsua. a b c d fig. 80 dyke completely converted into a metadyke, or amphibolite, (sample 1127). a, c: 1 nicol. b, d: x nicols. planar fabric mainly outlined by amphibole pleochroism. innermost saqqap kangerluarsua. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 68 of 96 geusbulletin.org subhorizontal axes (section 3.2). the effects on the dykes appear more penetrative. in addition to being folded together with the gneiss fabrics, the dykes north of saqqap kangerluarsua are completely converted into metadykes or amphibolites without any traces of the original igneous mineralogy or microstructure. however, the dykes do not lose their identity as dykes and are clearly recognisable as former intrusive bodies. the changes in both dyke and country rock mineralogy, from nuussuup qulaa to maligiaq, are summarised in figs 82 and 83. based on the mineral assemblages of metamorphosed dykes, the area between saqqap kangerluarsua and maligiaq can be divided into a number of metamorphic zones (figs 81, 84). 6.2.1 epidote zone the epidote zone extends from the southern shore of saqqap kangerluarsua to the south shore of ikertooq and akulleq (fig. 81). the zone is characterised by the common occurrence of epidote in quartzo-feldspathic gneisses and locally in metadykes. amphibole is the main ferro-magnesian mineral in metadykes, while garnet is an occasional additional phase. the original igneous clinopyroxene + plagioclase assemblage has been completely converted into a metamorphic amphibole + plagioclase ± garnet assemblage. these metadykes are proper amphibolites, consisting of about 60% amphibole, 30% plagioclase and 5% quartz (fig. 82). they have a granoblastic microstructure that usually has a well-developed planar and linear fabric (fig. 85), defined by a preferred orientation of elongated amphiboles. the amphiboles are blue-green, and plagioclases are typically oligoclases (fig. 82). biotite is a common minor constituent, and titanite and ilmenite may co-exist (fig. 85) with titanite often rimming ilmenite. compared to the almost completely metamorphosed dykes found locally on qaqqatoqaq, it is surprising that garnet is virtually itilleq amerloq avalleq m al ig ia q akulleq qeqertalik saqqap kangerluarsua qaqqatoqaq nuussuup qulaa inussuttusup tunua itillip ilua ikertooq sarfannguit n unaat 10 km amphibolite facies gneiss granulite facies gneiss charnockitic gneiss itillip ilua 66°45´ 66°30´53°00´ generally undeformed dykes generally undeformed dykes generally undeformed dykes northern limit of undeformed dykes cpx + grt in metadykes hbl + grt in metadykes epidote in qtz-fsp gneisses hbl + grt in metadykes opx in metadykes & qtz-fsp gneisses hbl in metadykes archaean pre-dyke granulite facies archaean pre-dyke amphibolite facies archaean pre-dyke amphibolite facies post(?) archaean pre-dykede formation and metamorphism archaean pre-dyke granulite facies sisimiut fig. 81 simplified geological map of the itillip ilua – ikertooq region, summarising metamorphic grades and mineral assemblages within both metadykes and host gneisses. see also fig. 84. dashed red line in centre of map marks the approximate southern boundary of pervasive nagssugtoqidian deformation and metamorphism. consequently, kangâmiut dykes north of this line are all converted to metadykes. south of the line partly deformed and metamorphosed dykes (margins or centres) occur, but these dykes are not completely deformed or metamorphosed. cpx: clinopyroxene. hbl: hornblende. grt: garnet. opx: orthopyroxene. qtz-fsp: quartzo-feldspathic gneiss. map based on fig. 2a in korstgård et al. 2006. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 69 of 96 geusbulletin.org absent within the epidote-zone metadykes. it has not been possible to find microstructural evidence of any garnet-consuming reactions and it is possible that – due to subtle bulk compositional differences – garnet was never produced in these metadykes. the typical mineral assemblage in the country gneisses is quartz + alkali feldspar + plagioclase + greenbrown biotite + epidote. other common minerals are blue-green amphibole and muscovite (fig. 83). the mineral assemblages in the quartzo-feldspathic gneisses of this epidote zone represent a downgrading or retrogression of the mineral assemblages seen in the archaean gneisses on qaqqatoqaq. amphibolites however, seem less affected, with clinopyroxene and garnet still being present. structurally the rocks in the epidote zone are characterised by intermediate (outcrop scale) to large-scale folding, around e–w-trending and subhorizontal axes, of both earlier archaean fabrics and hosted metadykes. no new fabrics are associated with these folds. further north, in avalleq, the nagssugtoqidian deformation becomes more intense. 6.2.2 clinopyroxene–garnet zone the disappearance of epidote and muscovite in the gneisses and an increased abundance of garnet in metadykes (fig. 86) mark an increase in metamorphic grade in an area that encompasses the outer part of avalleq and the southern part of akulleq (fig. 81). in the area between the south shore of ikertooq and halfway into maligiaq (fig. 81), clinopyroxene appears (fig. 87) and commonly occurs together with garnet, as well as amphibole and plagioclase in metadykes. the most common assemblage in this zone is plagioclase + amphibole + clinopyroxene ± garnet (fig. 88), whereas plagioclase + amphibole ± garnet is rare. titanite may dyke mineralogy in the itillip ilua–ikertooq region plagioclase titanite+opaques garnet orthopyroxene hornblende biotite quartz clinopyroxene epidote zone hbl zone cpx-grt zone opx zone 70 60 50 40 30 20 core rim plagioclase biotite hornblende titanite orthopyrox. clinopyrox. garnet s n ign ign ign+met ign+met met met brown green blue nuugssuup qula itillip ilua qaqqatoqaq saqqap kangerluarsua qeqertalik– avalleq avalleq– akulleq akulleqouter maligiaq inner maligiaq % a no rt hi te in p la gi oc la se ig ne ou s pl ag io cl as e m et am or ph ic p la gi oc la se fig. 82 variations in dyke and metadyke mineralogy across the itillip ilua – ikertooq region (location in fig. 2). upper right histograms show modal proportions of metamorphic minerals for metadyke samples north of qaqqatoqaq. in the main diagram, an content in plagioclase (rim to core) in dykes (black dots) and metadykes (blue dots) is shown. note systematic increase in an content in plagioclase in metadykes. note also the change in colour of amphibole, appearance of orthopyroxene and metamorphic clinopyroxene in northernmost zones. hbl: hornblende. cpx: clinopyroxene. grt: garnet. opx: orthopyroxene. ign: igneous. met: metamorphic. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 70 of 96 geusbulletin.org fig. 83 variations in country rock mineralogy across the itillip ilua – ikertooq region (location in fig. 2). there is a systematic northward change in colour of amphibole from blue-green to brown-green and in biotite from green-brown to red-brown for all rock types north of saqqap kangerluarsua. in quartzo-feldspathic gneisses orthopyroxene, clinopyroxene and garnet appear in northernmost zones, whereas epidote and muscovite disappear in these zones. bt: biotite. cpx: clinopyroxene. gr: graphite. grt: garnet. opx: orthopyroxene. pl: plagioclase. qtz: quartz. amphibolites micaschists (+qtz + pl + bt + gr) epidote zone cpx–grt zone opx zone orthopyroxene clinopyroxene garnet alkali-feldspar garnet sillimanite hornblende titanite orthopyroxene clinopyroxene garnet brown green blue (+/– qtz+ pl +/– bt) garnet-biotite gneisses (+qtz + pl + bt +/– gr) quartzo-feldspathic gneisses (+qtz + pl) alkali-feldspar biotite hornblende muscovite epidote orthopyroxene clinopyroxene garnet s n nuugsuup qula itillip ilua qaqqatoqaq saqqap kangerluarsua qeqertalik– avalleq avalleq– akulleq akulleq– outer maligiaq inner maligiaq brown green blue red brown green fig. 84 mineral parageneses across the ikertooq region from south (bottom of diagram) to north (top of diagram; see also fig. 81). the first two columns on the left are based on kangâmiut dyke mineralogy, the four columns to the right are based on host-rock mineralogy. the three rows at the bottom of the diagram show pre-nagssugtoqidian parageneses whereas the upper three rows show parageneses of nagssugtoqidian age. alk-fsp: alkaline feldspar. amp: amphibole. bt: biotite. cpx: clinopyroxene. ep: epidote. grt: garnet. opx: orthopyroxene. pl: plagioclase. qtz: quartz. sil: sillimanite. metamorphic zones de�ned by kangâmiut metadykes (nagssuqtoqidian) kangâmiut metadykes (pl + amp +/– qtz) metamorphic zones in host rocks (pre-nagssugtoqidian) amphibolites (pl + amp +/– qtz) garnet-biotite gneisses (qtz + pl + alk-fsp + bt) quartz0-feldspathic gneisses (qtz + pl +/– alk-fsp + bt) othopyroxe zone inner maligiaq clinopyrox.-garnet zone outer maligiaq epidote zone saqqap kangerluarsua– avalleq opx grt cpx + grt opx +cpx cpx + grt amp +/– grt +/– ep mostly igneous mineralogy. metamorphic overprinting. mostly igneous mineralogy. metamorphic overprinting. mostly igneous mineralogy. metamorphic overprinting. archaen granulite– amphibolite facies qaqqatoqaq pre-kangâmiut dyke ep-amphibolite facies itillip ilua archaen granulite– amphibolite facies nuussuup qulaa grt cpx + grt cpx + opx + grt grt cpx cpx + grt grt cpx +grt cpx cpx cpx + grt cpx cpx + grt cpx + opx grt grt + sil grt grt + sil no data grt + sil grt + sil no data amp opx opx + cpx amp + opx + cpx usually no alk-fsp amp grt ep ms ep + amp ep + ms amp amp + grt ep ep + amp amp + grt amp + opx + grt https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 71 of 96 geusbulletin.org fig. 85 metadyke in low amphibolite facies. a, b: foliated granoblastic microstructure with a planar fabric. c: amphibole and titanite (dirty brown). d: amphibole and biotite. a, c, d: 1 nicol. b: x nicols. sample 1138. north shore of inner saqqap kangerluarsua. a b c d fig. 86 metadyke in amphibolite facies. a, b: large plagioclase porphyroblast in amphibolite facies metadyke. c, d: with small garnet grains in amphibolite facies metadyke. a, c: 1 nicol. b, d: x nicols. sample 1055. head of avalleq. grt: garnet. grt a b c d https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 72 of 96 geusbulletin.org occur as an accessory phase but is partially replaced by ilmenite (fig. 89). the colour of amphibole has changed from grass-green to brown-green and the anorthite content of plagioclase has increased to well over 30% (fig. 82). epidote has disappeared from the host gneisses (fig. 83) and only traces of muscovite are present. the typical assemblage in the host gneisses is quartz + alkali feldspar + plagioclase + brown biotite, together with occasional green amphibole, clinopyroxene and garnet. 6.2.3 orthopyroxene zone the appearance of orthopyroxene in both metadykes and gneisses marks the transition from amphibolite to granulite facies. the orthopyroxene zone extends from halfway into maligiaq to the northernmost point mapped in fig. 81 and is characterised by the a b grt cpx fig. 88 metadyke from clinopyroxene–garnet zone. a: (1 nicol) garnet (grt; pink), clinopyroxene (cpx; light green), and amphibole (dark green). b: is the same with x nicols. sample 5257. north shore of inner akulleq. cpx a b c d fig. 87 upper amphibolite facies metadyke. a (1 nicol), b (x nicols): good foliation with amphibole (dark green) and clinopyroxene (light green). c (1 nicol), d (x nicols): clinopyroxene in centre. sample 5243. north shore of outer akulleq. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 73 of 96 geusbulletin.org occurrence of orthopyroxene in all rocks of suitable compositions. the most common assemblage in the metadykes is plagioclase + amphibole + orthopyroxene + clinopyroxene; less commonly plagioclase + amphibole + clinopyroxene + garnet (fig. 90). a single metadolerite with the anhydrous assemblage plagioclase + orthopyroxene + clinopyroxene is also found (fig. 91b). titanite is no longer present, amphibole is typically brown, and the anorthite content of plagioclase has further increased to around 40% (fig. 82). in quartzo-feldspathic gneisses, a typical mineral association is quartz + plagioclase + red-brown biotite + brown-green amphibole + orthopyroxene ± clinopyroxene (fig. 83). the almost complete absence of alkali feldspar is puzzling, since it cannot be attributed to any alkali-feldspar-consuming reaction in the gneisses. it is associated with a significant drop in the potassium fig. 89 titanite (ttn; centre of section) being replaced by ilmenite (ilm). a: 1 nicol. b: x nicols. sample 5237. north shore of inner akulleq. a b ilm ttn opx cpx a b c d fig. 90 granulite facies metadyke showing orthopyroxene (opx; pink), clinopyroxene (cpx; light green), amphibole (green) and ilmenite (opaque). a, c: 1 nicol. b, d: x nicols. sample 1195. inner maligiaq. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 74 of 96 geusbulletin.org content of the gneisses that must be either a primary feature or connected with metasomatic or partial melting processes (discussed further in chapter 7). the facies transition in maligiaq is very sharp and takes place within a few hundred metres in both dykes and gneisses. 6.2.4 summary of metamorphic mineral assemblages between saqqap kangerluarsua and maligiaq the progression of mineral assemblages during the metamorphism of basic rocks is typically illustrated by a series of acf diagrams (where a = al2o3 – k2o – na2o, c = cao, f = feo(tot) + mgo + mno, all molecular proportions). while most basic rocks are chemically complex and contain too many components to be accurately depicted in acf space (e.g. spear 1993), we find that the changes in mineral assemblages from saqqap kangerluarsua to maligiaq can be displayed schematically in acf diagrams. in fig. 92, we display actual mineral and whole-rock compositions from the three zones defined previously (sections 6.2.1–6.2.3). for the two northern zones, the lowest variance assemblage observed was chosen to represent the metamorphic conditions and mineral diversity. for the southernmost zone (epidote zone) two different mineral assemblages are shown (samples 1132 and 1128). note that due to limitations mentioned in previous sections, crossing tie-lines (e.g. plagioclase–orthopyroxene crossing garnet–clinopyroxene) does not necessarily mean that these mineral reactions took place (spear 1993). in the epidote zone, between saqqap kangerluarsua and avalleq, mineral assemblages are plagioclase + amphibole ± garnet ± epidote. generally, garnet occurs in more fe-rich bulk compositions, whereas epidote is more common in ca-richer rocks. in one sample, however (1132, shown in fig. 92), all four phases are present. this marks the transition to lower amphibolite facies (e.g. bucher & frei 1994). in the clinopyroxene–garnet zone, between avalleq and akulleq, all epidote has been consumed, and the typical assemblage is plagioclase + amphibole + garnet, with garnet becoming volumetrically more important than in the lower grade zone (no samples from this zone were analysed). clinopyroxene appears between avalleq and outer maligiaq, and a common assemblage is now plagioclase + amphibole + clinopyroxene + garnet (represented by sample 5262 in fig. 92). this assemblage is typical of upper amphibolite facies. granulite facies conditions are reached in the orthopyroxene zone, in the inner part of maligiaq. the mineral assemblages are the same as in the clinopyroxene–garnet zone, except that orthopyroxene is now stable in rocks of suitable composition. orthopyroxene-bearing and orthopyroxene-free assemblages can coexist at the same temperature, if the proportion of water in the attending fluid phase varies from low to high, respectively. in this zone, the general assemblages found are plagioclase + amphibole + clinopyroxene ± orthopyroxene ± garnet (all phases present in sample 1195 in fig. 92). figure 82 shows the appearance and modal increase in both orthopyroxene and clinopyroxene in the orthopyroxene zone. throughout this sequence, amphiboles are increasingly consumed through dehydration reactions (decreasing modal contents; figs 50, 82), and their compositions are continuously changing (migrating from close to the ‘c–f’ line towards the ‘a’ apex in acf diagrams; fig. 92). in addition to pressure and temperature, the nature and progress of these amphibole-consuming reactions is a function of bulk-rock composition and fluid composition. however, in the present sequence of rocks, the ultimate disappearance of amphiboles is never reached. metamorphic plagioclase similarly changes composition as a result of the reactions fig. 91 granulite facies metadyke. a: coexisting orthopyroxene, clinopyroxene, amphibole, garnet and plagioclase. 1 nicol. sample 5203. b: coexisting orthopyroxene, clinopyroxene and plagioclase. 1 nicol. sample 5248. inner maligiaq. amp: amphibole. cpx: clinopyroxene. grt: garnet. opx: orthopyroxene. b cpx amp opx grt opx cpx a https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 75 of 96 geusbulletin.org that characterise the progression from saqqap kangerluarsua to maligiaq. while the acf projections in fig. 92 do not show compositional variations of plagioclase, fig. 82 shows how the anorthite content of plagioclase increases from itillip ilua to maligiaq. 6.3 types of facies transitions in the itillip ilua – ikertooq region 6.3.1 facies transitions within the itillip ilua – ikertooq region, three types of facies transitions or boundaries are recognised, based on whether accompanying metamorphism was prograde or retrograde as well as on the intensity of the accompanying deformation (korstgård 1979b). the first two pre-date kangâmiut dyke emplacement, whereas the last one affects the dykes as well as their host rocks. 1. the amphibolite-granulite facies transition in the archaean areas of qaqqatoqaq and nuussuup qulaa (fig. 93) is prograde and static, in the sense that the boundary was not established as a result of a deformational event but reflects static equilibration of mineral assemblages to the conditions that prevailed when the rocks were at their deepest crustal level. during later uplift, the rocks escaped any significant metamorphic changes (retrogression) due to absence of deformation. 2. the granulite to low amphibolite facies and amphibolite to low amphibolite facies transitions along itillip ilua are retrograde and dynamic, in the sense that these were established as a direct consequence of the shearing along the itilleq shear zone. mineral assemblages in the shear zone equilibrated to the metamorphic conditions at a higher crustal level, while the shearing facilitated the retrogression. 3. the amphibolite–granulite facies transition in maligiaq (fig. 93) is both prograde and dynamic. it could be considered a displaced prograde and static transition that has been thrusted into a subvertical position. its position at the present level of erosion is clearly a result of deformation. the nature of thrusting is that rocks from a deeper crustal level are brought up on the hanging-wall side and a greater load is placed on rocks on the foot-wall side. due to this, regions bounding a ductile overthrusting shear zone will suffer metamorphic and deformational effects not experienced by regions adjacent to ductile transcurrent shear zones, which are largely unaffected by the shearing. 6.3.2 metamorphic boundaries and magnetic anomalies a magnetic anomaly map for the ikertooq region shows a remarkable correlation with metamorphic boundaries (fig. 94; korstgård et al. 2006). thus, within the archaean granulite facies areas, just north and south of itillip ilua (‘a’ in fig. 94), strong magnetisations fig. 92 data from four representative dyke samples between saqqap kangerluarsua and maligiaq. acf diagrams include sample numbers and illustrate progressive changes in mineral assemblages from lower amphibolite to granulite facies. mineral analyses for plagioclase, clinopyroxene, amphibole, garnet, and orthopyroxene are plotted under the assumption that all fe is fe2+, whereas all fe in epidote (sample 1132) is assumed to be fe3+. for consistency, whole-rock data are also plotted assuming all fe to be fe2+. in all mineral assemblages, quartz and a hydrous fluid phase are assumed present. a: (al2o3 – k2o – na2o). c: cao. f: (feo(tot) + mgo + mno), all molecular proportions. amp: amphibole. cpx: clinopyroxene. ep: epidote. grt: garnet. opx: orthopyroxene. pl: plagioclase. wr: whole-rock data. qeqertalik sallersua sarfa nnguit n unaat amerloq sisimiut ikertooq m al ig ia q akulleq avalleq saqqap kangerluarsua 66°45´ 53°00´ 10 km n "charnockitic" gneiss granulite facies gneiss amphibolite facies gneiss supracrustal gneiss amphibolite dyke metadyke a c f 1195 a c f 1128 a c f 5262 a c f 1132 pl grt cpx opx amp wr ep https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 76 of 96 geusbulletin.org are attributed to a high content of magnetite, formed through the breakdown of hydrous mafic minerals (biotite, amphibole) during the transition from amphibolite facies to granulite facies. the gradual increase in magnetic intensity, from b to c to a in fig. 94, even conforms to this gradual prograde facies transition. the e–w-trending low magnetic anomaly, coincident with the itilleq shear zone (‘d’ in fig. 94), is most likely caused by breakdown of magnetic minerals during metamorphic retrogression, from granulite facies or upper amphibolite facies (‘a’, ‘c’ and ‘b’ in fig. 94) to epidote amphibolite facies in the shear zone (‘d’ in fig. 94). the post-dyke lower amphibolite facies rocks along qeqertalik (‘e’ and ‘f’ in fig. 94) also display low magnetic anomalies, whereas the upper amphibolite to granulite facies rocks along the ikertooq shear zone and north display high magnetic anomalies (‘h’ and ‘g’ in fig. 94). isolated high intensity anomalies can also locally be correlated with distinct lithologies or intrusions (e.g. the anorthosite complex at ‘j’ in fig. 94). the presence or absence of kangâmiut dykes is not reflected in the aeromagnetic data. the observed correlation between metamorphic facies, deformation and magnetisation can be extended to other areas of the southern nagssugtoqidian orogen (fig. 95). sisimiut ikertooq qeqertalik saqqap kangerluarsua nuussuup qulaaitillip ilua ikertooq m al ig ia q post-dyke (nagssutoqidian) granulite facies post-dyke (nagssugtoqidian) amphibolite facies pre-dyke (archaean) amphibolite facies pre-dyke amphibolite facies pre-dyke (archaean) amphibolite facies pre-dyke granulite facies pre-dyke (archaean) granulite facies mainly undeformed dykes only deformed dykes 10km 53°00ʹ 66°45ʹ 66°30ʹ53°00ʹ66°30ʹ 66°45ʹ fig. 93 simplified map of the itillip ilua – ikertooq area, showing the distribution of relative preversus post-dyke ages of the main metamorphic facies, south and north of a boundary (red dashed line) through saqqap kangerluarsua. thick black lines: undeformed or only partly deformed dykes. thick dashed green lines: deformed dykes (only a few shown). thin black dashed lines: planar structures in host rocks. the dashed red line marks the southern limit of completely deformed and metamorphosed kangâmiut dykes. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 77 of 96 geusbulletin.org fig. 94 magnetic anomaly map for the itillip ilua – ikertooq area (modified from korstgård et al. 2006) showing the strong correlation between metamorphic boundaries and the magnetic signatures. a–j: magnetic anomaly centres described in korstgård et al. 2006. the dashed white lines correspond to the metamorphic facies boundaries in fig. 93. within the archaean granulite facies areas, just north and south of itillip ilua (a), strong magnetisations are attributed to a high content of magnetite, formed through the breakdown of hydrous mafic minerals (biotite, amphibole) during the transition from amphibolite to granulite facies. the gradual increase in magnetic intensity, from b to c to a, even conforms to this gradual prograde facies transition. the e–w-trending low magnetic anomaly, coincident with the itilleq shear zone (d), is most likely caused by breakdown of magnetic minerals during metamorphic retrogression, from granulite facies or upper amphibolite facies (a, b, c) to epidote amphibolite facies in the shear zone (d). the post-dyke lower amphibolite facies rocks along qeqertalik (e and f) also display low magnetic anomalies, whereas the upper amphibolite to granulite facies rocks along the ikertooq shear zone and northwards display high magnetic anomalies (h and g). isolated high intensity anomalies can also locally be correlated with distinct lithologies or intrusions (e.g. the anorthosite complex at j). [nt] 526 424 365 319 282 247 216 189 164 140 1189673523211–7 –26 –46 –68 –89 –110 –131 –150 –168 –185 –202 –222 –245 –266 –283 –305 –327 –352 –377 –398 –409 –436 52°30'53°30' 66 °5 0' 66 °4 0' 66 °3 0' 53° a b c d dd e e f f g g h h i 10 km j sisimiut qeqertalik itillip ilua ikertooq m aligaaq saqqap kangerluarsua 50°52° 66 °4 0' 67 °0 5' 66 °1 5' fig. 94 53° kangerl ussu aq sisimiut kangerlussuaq saqqap kangerluarsua itillip ilua ikertooq qeqertalik [nt] ?? 50 km 536 407 334 284 243 209 178 151 125 76 54 32 11 –9 –30 –50 –72 –91 –110 –127 –145 –161 –176 –190 –204 –219 –234 –264 –282 –303 –326 –354 –386 –416 –461 –535 fig. 95 total intensity magnetic field anomaly map across the southern nagssugtoqidian orogen. the southern foreland of the nagssugtoqidian orogen from the coast to the greenland ice sheet is shown along with the location of the itillip ilua – ikertooq region. white frame: locates fig. 94. ?: location of thrust uncertain. (modified from korstgård et al. 2006). https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 78 of 96 geusbulletin.org 7 chemical changes in dykes and country rocks from nagssugtoqidian deformation and metamorphism the kangâmiut dyke swarm was variably affected by metamorphism and deformation from area 01 to area 11 (see chapters 5 and 6), but as shown in section 4.2, the dominant compositional variations observed were those associated with magmatic differentiation processes – mainly fractional crystallisation and minor crustal contamination (mayborn et al. 2008) – with expected differences between chilled margins and later, more leucocratic interiors. there is no evidence that the metamorphic and structural overprints on the dykes in areas 01 to 07 significantly altered their bulk composition. however, some data are suggestive of potential element mobility in areas farther north. to investigate the nature and magnitude of any bulk chemical changes that the dykes and their country rocks may have experienced during amphibolite and granulite facies metamorphism in the northernmost areas, we will look at more mobile large-ion lithophile (lil) elements identified earlier, and major and trace elements traditionally associated with increased mobility during medium to high-grade metamorphism (e.g. bridgwater 1979, rollinson 1993). the purpose of identifying and quantifying any compositional variations due to deformation and metamorphism is to ensure that observed metamorphic mineral assemblages reflect metamorphic conditions and not a combination of metamorphism and bulk compositional changes. 7.1 kangâmiut dykes figure 96 displays the south-to-north (area 01 to area 11, see fig. 46) compositional variations for selected elements of interest. to minimise any effects of fractionation, we have normalised each of the potentially mobile elements to an incompatible element (zr). the two most mobile lil elements (rb, k) show minimal or non-systematic variations between areas 01 and 08, a slight increase from area 08 to 09 and then a marked decrease into areas 10 and 11. both k and rb decrease from area 09 to 11, however, rb decreases more (cf. fig. 56a), so k2o/rb increases. other lil elements (ba, sr, eu) do not show any significant variation (shown in fig. 56a). pb and u both decrease from area 09 to 11, in accordance with behaviour typically observed during regional amphibolite and granulite facies metamorphism (e.g. bridgwater 1979). both sm and nd display similar, albeit muted, trends (fig. 56a). the ‘pseudo’-trends seen in the mgo vs k2o plot (fig. 53) are, in light of the above, interpreted to represent normal variations to be expected among samples fig. 96 variations in dyke chemistry from south (area 01, left) to north (area 11, right) for those elements that display significant changes along the transect. to minimise the effect of fractionation, k2o and rb have been normalised against an immobile element, zr (ppm). each coloured circle represents an analysis. size of bins along y-axis: 1/30 of entire data range. the red diamonds show the mean value within each area. rb decreases in areas 10 and 11, whereas k2o only shows a moderate decrease, so the ratio between these elements shows a significant increase in the northernmost areas. h2o (loi: loss on ignition, in wt%) also shows lower values in areas 10 and 11, in accordance with dehydration reactions taking place from amphibolite to granulite facies. 0.01 0.1 1 10 0.01 0.1 1 0.001 0.01 0.1 area 01 area 02 area 03 area 04 area 05 area 06 area 07 area 08 area 09 area 10 area 11 0 1 2 3 h2o(loi) (k2o/zr)/(rb/zr) k2o/zr rb/zr https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 79 of 96 geusbulletin.org taken from margins and centres and from the leucocratic patches that are found in some dykes. the k and rb patterns in areas 08 to 11 are best explained by the gradual breakdown of biotites and amphiboles with increasing metamorphic grade. the incompatible elements k and rb are preferentially partitioned into the aqueous phase generated during breakdown of the hydrous minerals. thus the granulite facies dykes are gradually depleted of these elements. the gradual breakdown was also demonstrated in fig. 50, where a northwards volumetric decrease in hydrous minerals was evident. the decrease in the volume of hydrous minerals is also shown by the overall decrease in analysed h2o (by loss on ignition; loi). in fig. 96, the range of h2o (loi) is shown for all areas, and the decrease in areas 09–11 relative to the remainder is clear. within individual dykes south of area 09, there are similarly no signs of preferential lil element mobility between margins and centres. figure 97 shows available lil data (k, rb) from samples taken from margins to centres of dykes from areas 02 to 08. to minimise the effect of fractionation, each element has been normalised to zr. in this diagram, each block represents data collected from margins (left) to centres (right) of individual dykes. we see that k2o tracks rb consistently, with rb showing larger variations. 7.2 quartzo-feldspathic gneisses, amphibolites and supracrustal rocks data from the host rocks of the dykes, including tonalitic to granodioritic gneisses, amphibolites and various rocks of supracrustal origin, were collected in areas 07–11 with the aim of characterising these lithologies and assessing any compositional changes associated with the increasing metamorphic grade (korstgård 1980). in section 7.1, we documented the chemical changes associated with element mobility driven by high-grade metamorphism in the kangâmiut dykes. here we will briefly see if similar changes are recorded by the host rocks. the only elements that show a systematic and significant variation from areas 07 to 11 are k and rb (fig. 98). similar to the other element mobility plots, we have normalised the data to an incompatible element (zr). the gneisses show k2o and rb decreasing from area 09 to 11, and since the rb decrease is comparatively larger, k2o/rb increases. this was the same trend observed in the kangâmiut dykes, hosted by these gneisses. similar trends may be hinted at by the amphibolites (one sample) and supracrustal gneisses (two samples), but the limited number of samples in these units from area 11 precludes a definitive determination. area 02 area 03 area 04 area 05 06 07 08 1 0.1 0.01 0.001 rb/zr k2o/zr 40 m 65 m 60 m 40 m 60 m 50 m 100 m 45 m 120 m 60 m 30 m h w h h h h b bb mh h h h h h k fig. 97 mobility of available lil elements from margin to centre in individual kangâmiut dykes in areas 02 to 08. data from h: hansen (1989). w: windley (1970). b: bridgwater (unpublished data). m: mayborn (2000). k: korstgård (1980). each block of data represents a single dyke sampled from margin (left) to centre (right). most dykes show normalised k2o (green) tracking rb (red) with no systematic changes. consequently there are no signs of k2o:rb increasing anywhere in the dyke in contrast to that observed in the higher-grade metamorphic zones (area 10 and 11). k2o in wt%, rb and zr in ppm. dyke orientation (thick black lines) and thickness data (in m) are shown where available. ssw nne 0.01 0.1 1 0.01 0.1 1 10 0.0001 0.001 0.01 0.1 k 2o /z r rb /z r (k 2o /z r) /( rb /z r) 0 250 500 750 amphibolitesqtz-fsp gneisses supracrustals area 07 area 08 area 09 area 10 area 11 (maligiaq)(itillip ilua) fig. 98 variations in k2o and rb in rocks hosting kangâmiut dykes from areas 07 to 11. data from quartzo-feldspathic gneisses, amphibolites and rocks of supracrustal origin are projected on to a c. 60 km long nne– ssw-trending line from maligiaq (right) to itillip ilua (left; see fig. 2 for location of projection line; each unit along x-axis is c. 50 m). k2o and rb are normalised against an immobile element, zr. all lithologies show unsystematic and quite scattered variations in areas 07 and 08. from area 09 to 11, quartzo-feldspathic gneisses show a decrease in both k2o and rb, however since the rb decrease is larger, the resultant k2o:rb increases dramatically in area 11. amphibolites and supracrustal gneisses show vaguely similar trends but those are based on relatively few samples. the smooth trend lines are calculated with the locally weighted least squares regression method. k2o in wt%, rb and zr in ppm. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 80 of 96 geusbulletin.org 8 assessment of metamorphic conditions from observed mineral reactions and equilibria establishing the pressure and temperature conditions that the kangâmiut dykes have experienced can provide useful constraints for reconstructing their tectonic history. although numerous approaches are available for geothermobarometric calculations, a common underlying assumption for all is that thermodynamic equilibrium has been achieved among the mineral phases of interest. it is clear, however, from the evidence presented in previous chapters, that the recrystallisation history for these rocks has been complex and episodic. preservation of compositional zoning, symplectites, mineral overgrowths and reaction rims attest to localised compositional domains. additionally, deformation may result in the juxtapositioning of minerals that had not thermodynamically equilibrated. consequently, although the following discussion presents the quantitative results from rigorously applied geothermobarometric calculations, subjective decisions played a role in selecting which samples would be used and how results are interpreted. these choices are discussed next. although current approaches for calculating pseudosections (e.g. forshaw et al. 2018) are powerful tools for establishing p–t conditions for metamorphic rocks, these also rely on the assumption that thermodynamic equilibrium has been achieved for the selected bulk-rock composition used in the calculation. from our discussions of mineralogy and textures, it is obvious that equilibrium is not satisfied for many of the samples that we have data for. therefore, although we have pursued developments in such analyses (j. forshaw, personal communication 2021), the results, as expected, are inconsistent with the microanalytical results we have obtained and are not shown here. 8.1 geothermometry 8.1.1 methodology a suite of geothermometers was used for samples with appropriate mineralogies (table 2). clinopyroxene–orthopyroxene brey et al. (1990) and clinopyroxene–garnet (ellis & green 1979) temperatures were often consistent with those derived from other geothermometers. however, these geothermometers were relevant for only a few samples, which made the amphibole–plagioclase (blundy & holland 1990) and ti-in-amphibole (liao et al. 2021) formulations more useful. in addition, separating remnant igneous pyroxene compositions from metamorphic compositions was not always possible. therefore, in the following discussion we focus on results from the two amphibole-dependent formulations. in those instances where garnet and clinopyroxene were in contact, we used the ellis & green (1979) method to calculate temperatures, which generally provided results similar to those obtained from the blundy & holland (1990) and liao et al. (2021; table 3) formulations. the ti-in-amphibole geothermometer (tiat; liao et al. 2021) is calibrated for ca-bearing amphiboles and is pressure-independent. it relies solely on the ti content in the tetrahedral coordination of amphiboles. as such, it is easy to use. however, because a range of amphibole compositions can form under different temperature and pressure conditions, not all amphibole analyses from a given sample can be used to establish temperature conditions. as a result, analyses from the 53 samples used for this calculation were screened to assure that they met the compositional restrictions defined by liao et al. (2021). the amphibole-plagioclase geothermometer (apt) of blundy & holland (1990) is based on cation abundances in an amphibole’s a and t1 sites and therefore depend on appropriate site occupancy assignments. the 2020 updated spreadsheet by locock (2014) was used to assign cation site occupancies in the analysed amphiboles. the computed temperature is also dependent upon the anorthite content of coexisting plagioclase. given the complex history of these rocks and the significant degree of igneous compositional zoning within these plagioclases, it is obviously problematic to identify the appropriate amphibole and plagioclase composition to use in these calculations. adding to this complexity is the fact that the derived geothermometer is pressure dependent. we have therefore chosen to calculate temperatures at 8 kbar (kilobar) from randomly paired plagioclase and amphibole analyses in each sample and identified temperature clusters within these results. these clusters consistently ranged from approximately 550°c to 700°c (“low” group in table 3), 700°c to 800°c (“med.” group in table 3) or were greater than 800°c (“high” group in table 3). then, to avoid any uncertainties relating to pressure dependencies, we averaged tiat temperatures for the amphiboles inside table 2 geothermometers and geobarometers used in this study mineral assemblage parameter number of samples reference clinopyroxene–orthopyroxene t 12 brey et al. (1990) clinopyroxene–garnet t 20 ellis & green (1979) amphibole–plagioclase t 53 blundy & holland (1990) ti in ca-amphibole t 53 liao et al. (2021) garnet–orthopyroxene p 9 brey et al. (1990) amphibole–garnet p 38 kohn & spear (1990) https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 81 of 96 geusbulletin.org ta bl e 3 pr es su re –t em pe ra tu re (p –t ) d at a us ed in th is s tu dy . sa m pl e no . u tm e u tm n am p– pl am p– pl am p– pl am p– pl ti -in -a m p ti -in -a m p ti -in -a m p ti -in -a m p ti -in -a m p g rt –o px cp x– o px g rt –c px am p– g rt t: b & h p: 8 k ba r (a ve .) t: b & h p: 8 k ba r (lo w ) t: b & h p: 8 k ba r (m ed .) t: b & h p: 8 k ba r (h ig h) t: l (a ve .) t: l (s d ) t: l p: b & h (lo w t ) t: l p: b & h (m ed . t ) t: l p: b & h (h ig h t) p: b , k & n t: b , k & n t: e & g t: l p: k & s (° c) (° c) (° c) (° c) (° c) (° c) (k ba r) (k ba r) (k ba r) (k ba r) (° c) (° c) (k ba r) 10 11 41 12 55 74 04 76 7 77 1 75 8 73 1 16 9. 7 10 40 43 80 47 74 14 09 0 70 6 69 5 74 9 67 9 27 9. 0 12 .4 10 87 42 10 68 73 99 13 6 76 8 75 6 65 5 12 14 .3 5. 15 10 95 42 46 17 73 99 10 4 72 8 62 8 72 2 73 8 56 6. 8 64 6 9. 30 11 06 42 20 68 74 00 29 0 73 8 73 6 78 4 3 5. 0 9. 50 11 16 42 20 68 74 00 29 0 73 7 72 8 69 7 21 10 .0 11 28 42 96 11 74 02 54 7 73 7 74 1 80 8 16 3. 7 11 .3 6 11 32 42 94 60 74 03 18 0 74 3 65 5 72 6 84 9 74 8 19 6. 5 12 .5 7. 11 11 45 43 26 10 74 25 87 5 78 6 76 3 83 5 20 3. 5 69 8 9. 09 11 95 43 64 87 74 28 72 1 82 6 80 2 80 7 12 5. 8 6. 3 71 5 12 85 42 79 46 73 90 02 0 77 8 77 7 84 1 77 4 12 8. 6 11 .9 28 78 5 83 8 7. 82 12 86 42 78 95 73 90 52 9 72 6 66 4 73 0 78 3 17 4. 6 12 94 41 87 47 73 91 45 7 79 7 77 0 82 5 74 7 36 8. 6 11 .8 73 5 7. 20 12 97 41 73 38 73 85 65 5 76 2 75 5 78 9 7 5. 9 12 .6 5 13 44 41 42 56 73 81 33 0 76 4 56 1 75 3 87 0 71 9 10 .1 16 .2 23 76 8 63 0 7. 27 13 50 41 58 15 73 96 74 9 71 2 71 0 84 3 21 65 0 11 .5 0 52 62 44 29 82 74 15 14 2 81 6 78 2 81 5 72 8 15 11 .7 12 .1 73 6 6. 95 52 68 44 83 10 74 14 71 3 69 5 69 0 73 9 72 0 30 6. 3 9. 5 87 77 7 40 58 77 72 82 88 9 74 2 67 9 74 2 80 2 60 2 45 10 .9 7. 63 15 80 27 38 28 75 73 69 45 1 87 5 83 5 78 6 18 10 .8 15 80 35 39 25 32 73 03 37 5 84 4 59 6 78 9 91 5 61 8 5 9. 8 8. 0 9. 9 14 10 80 41 43 02 38 97 92 73 90 19 3 76 4 76 3 82 5 82 2 33 4. 2 8. 2 20 78 0 9. 16 41 43 04 39 44 75 73 84 98 0 90 7 91 0 82 7 33 12 .3 87 5 41 43 98 41 88 94 73 84 77 1 80 3 67 4 77 4 76 6 40 8. 5 78 8 7. 41 41 43 99 41 88 94 73 84 77 1 72 5 65 8 72 5 82 0 70 7 37 4. 6 9. 1 14 .7 62 5 8. 88 41 44 04 41 55 89 73 85 82 7 72 0 65 8 74 0 75 6 34 3. 3 7. 5 62 5 9. 05 41 44 07 41 88 94 73 84 77 1 74 2 74 1 76 4 14 6. 5 73 4 41 44 09 41 07 11 73 83 88 6 76 7 61 8 76 3 77 5 34 8. 0 64 7 8. 16 41 44 10 41 07 11 73 83 88 6 70 2 57 1 72 3 73 0 12 7. 5 67 3 9. 08 41 44 11 41 07 11 73 83 88 6 78 2 74 3 80 8 35 4. 4 11 .9 0 41 44 27 41 55 89 73 85 82 7 69 9 64 5 79 5 18 10 .3 7 41 44 29 41 55 89 73 85 82 7 56 1 59 1 75 0 76 3 25 7. 3 3 82 5 7. 66 41 44 32 39 24 38 73 85 11 2 78 8 75 8 78 0 41 6. 6 73 3 13 .2 1 41 44 33 39 10 89 73 87 33 5 77 8 76 9 84 6 80 9 13 5. 5 10 .1 73 5 7. 35 41 44 34 42 68 07 74 13 31 7 70 3 70 9 77 8 36 3. 5 9. 37 https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 82 of 96 geusbulletin.org ta bl e 3 (c on tin ue d) p re ss ur e– te m pe ra tu re (p –t ) d at a us ed in th is s tu dy . sa m pl e no . u tm e u tm n am p– pl am p– pl am p– pl am p– pl ti -in -a m p ti -in -a m p ti -in -a m p ti -in -a m p ti -in -a m p g rt –o px cp x– o px g rt –c px am p– g rt t: b & h p: 8 k ba r (a ve .) t: b & h p: 8 k ba r (lo w ) t: b & h p: 8 k ba r (m ed .) t: b & h p: 8 k ba r (h ig h) t: l (a ve .) t: l (s d ) t: l p: b & h (lo w t ) t: l p: b & h (m ed . t ) t: l p: b & h (h ig h t) p: b , k & n t: b , k & n t: e & g t: l p: k & s 41 44 38 41 33 66 73 97 70 7 77 3 74 6 82 5 76 4 24 6. 8 11 .6 7 61 5 7. 86 41 44 42 39 30 73 73 80 27 1 67 9 66 8 75 7 7. 69 41 44 43 39 30 73 73 80 27 1 81 3 64 4 73 9 82 9 76 3 44 6. 5 11 .9 7. 69 41 44 45 39 37 87 73 79 66 3 77 2 67 7 74 7 68 7 23 7. 3 11 .8 53 0 7. 95 41 44 51 43 31 89 74 27 18 6 79 4 78 8 83 0 85 3 18 4. 5 6. 7 95 9 9. 20 41 44 76 43 53 69 74 28 61 8 77 8 76 5 82 8 82 5 9 4. 3 8. 2 31 71 4 72 2 8. 44 41 44 79 43 26 21 74 27 56 8 72 0 60 9 73 0 76 4 56 2. 7 5. 8 22 74 3 58 6 43 02 67 39 90 57 72 94 06 1 73 3 62 3 76 6 80 4 81 6 27 5. 1 7. 2 84 8 43 02 74 39 90 57 72 94 06 1 76 9 60 3 78 4 88 1 59 6 37 80 1 43 02 75 39 90 57 72 94 06 1 83 8 63 8 76 2 91 2 73 3 9. 7 17 .5 13 81 0 7. 63 43 02 85 39 90 57 72 94 06 1 79 4 76 9 83 0 61 1 13 .5 5 43 02 88 38 59 57 73 14 30 3 77 2 57 5 76 0 86 9 73 8 46 9. 3 15 .3 43 02 89 38 59 57 73 14 30 3 95 7 73 5 95 3 33 43 02 90 38 59 57 73 14 30 3 74 4 74 3 57 5 18 .6 13 .0 9 43 02 94 38 59 57 73 14 30 3 73 9 64 7 74 7 60 8 38 16 .8 10 .4 9 43 02 95 38 59 57 73 14 30 3 77 4 69 3 72 7 59 5 13 .4 0 43 03 03 39 23 60 73 23 94 3 81 7 60 4 85 3 79 7 80 11 .1 7. 19 ab br ev ia tio ns : a m p: a m ph ib ol e. c px : c lin op yr ox en e. e p: e pi do te . g rt : g ar ne t. o px : o rt ho py ro xe ne . p l: pl ag io cl as e. t i: tit an iu m . u tm : u ni ve rs al t ra ns ve rs e m er ca to r co or di na te s. c al ib ra tio ns u se d: b & h : b lu nd y & h ol la nd (1 99 0) . l : l ia o et a l. (2 02 1) . e & g : e lli s & g re en (1 97 9) . b , k & n : b re y et a l. (1 99 0) . k & s: k oh n & s pe ar (1 99 0) . https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 83 of 96 geusbulletin.org each data cluster and used that temperature for our pressure calculations described below. this approach generally provided broadly consistent results for geographically neighbouring samples, as well as allowed for a screening of compositional outliers. 8.1.2 results the computed temperatures range from 550°c to c. 900°c (table 3). temperatures greater than 850°c are most commonly associated with samples that retain igneous textures. we interpret these samples to reflect partial re-equilibration of a remnant igneous mineralogy. lower temperatures are commonly associated with clear textural evidence of metamorphic recrystallisation, as discussed earlier (chapter 5). the regional pattern of temperatures is shown in fig. 99. while this pattern is still complex, there are several features that stand out. in the vicinity of the maligiaq, where the ikertooq shear zone is well-defined, samples north of and inside this shear zone consistently preserve temperatures that exceed 800°c. immediately south of that shear zone, temperatures are consistently lower; commonly reaching values in the range of 700°c to 770°c. these results are consistent with an upper amphibolite facies – granulite facies boundary that marks a transition from a lower grade footwall block and into an over-thrusted higher 50 km sisimiut charnockite variably reworked archean gneisses supracrustal rocks ma�c dykes kangâmiut ma�c dyke swarm undi�erentiated archean/paleoproterozoic paleoproterozoic ikertooq sisimiut sisimiut isortuat kangaamiut sermiat n fitillip ilua kangaamiut maniitsoq so ut he rn n ag ss ug to qi di an fo re la nd nassuttooq kangerl ussu aq nordre isortoq shear zone ikertooq shear zone s s c n o sn o isortoq kangerlussu atsia q greenland ice sheet 52° 67° 67° 66° 54° 52°54° d k d southern limit to >820°c northern limit to <700°c t between 550°c & 700°c t between 700°c & 830°c fig. 99 distribution of samples for which temperature calculations are available (see table 3). in the legend, t is calculated temperatures. white circles overlapping red circles indicate samples in which temperatures formed distinct clusters in both groups. the red line (dashed where data are insufficient to tightly constrain location) indicates the southern limit of temperatures that exceeded 820°c; the blue dashed line is the northern limit of samples with temperatures below 700°c. data sources: d: davidson (1979). k: preliminary results from kriegsman et al. (1996). other symbols and abbreviations in fig. 1. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 84 of 96 geusbulletin.org grade hanging wall block, as described in chapter 6. samples scattered throughout the central and southern areas record temperatures between 550°c and 700°c. the boundary north of which temperatures are consistently above 700°c is marked in fig. 99. these lower temperatures may record late-stage retrogressions. 8.2 geobaromometry 8.2.1 methodology pressure was calculated using several different approaches. the amphibole–garnet geobarometer (agp) of kohn & spear (1990) was used when these two phases were present. the amphibole cation site occupancy used in this calculation was the same as the one used for apt calculations. garnet end member proportions for grossularite, pyrope, almandine and spessartine were determined using the stipulations by kohn & spear (1990). temperatures used in the calculations were those derived in section 8.1.2 (table 3; fig. 99). the garnet–orthopyroxene barometer of brey et al. (1990) was also used, when these mineral pairs coexisted (table 3). finally, we used an approach that combines the ti-in-amphibole temperature, following liao et al. (2021), with pressure-dependent temperatures that were computed from coexisting amphiboles and plagioclases, as prescribed by blundy & holland (1990). in this approach, we determined the pressure that would give the same temperature as the ti method by liao et al. (2021). all results are listed in table 3. since the latter method gave results for most samples, it is these results that we used in mapping the pressure variability across the region (fig. 100). fig. 100 distribution of metamorphic pressures (in kbar) between maniitsoq and sisimiut. grey circles: locations of samples used to constrain the isobars. ??: indicates uncertainty in the location and trend for the 7.5 kbar boundary. other symbols and abbreviations in fig. 1. sisimiut charnockite variably reworked archean gneisses supracrustal rocks ma�c dykes kangâmiut ma�c dyke swarm s undi�erentiated archean/paleoproterozoic paleoproterozoic n f so ut he rn n ag ss ug to qi di an fo re la nd nordre isortoq shear zone ikertooq shear zone s c n o sn o greenland ice sheet ikertooq sisimiut sisimiut isortuat kangaamiut sermiat itillip ilua kangaamiut maniitsoq nassuttooq kangerl ussu aq isortoq kangerlussu atsia q 52° 67° 67° 66° 54° 52°54° 10.0 kbar 9.0 kbar 7.5 kbar 7.5 kbar 7.5 kbar 9.0 kbar 10.0 kbar ?? ?? https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 85 of 96 geusbulletin.org 8.2.2 results the computed pressures range from c. 5 kbar to c. 13 kbar (table 3). the higher and lower values in this broad range were based on amphibole–garnet pairs and could not be repeated with data from amphibole–plagioclase pairs from adjacent samples. we conclude that the amphiboles and garnets were likely not in thermodynamic equilibrium and thus do not represent actual p–t conditions experienced by samples. pressure estimates based on coexisting amphibole and plagioclase showed much narrower intraand inter-sample variations, and the isobars contoured in fig. 100 were constructed from this perspective. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 86 of 96 geusbulletin.org 9 discussion in this chapter, we briefly summarise the key findings and main conclusions and attempt to fit them into a coherent chronological and tectonic framework describing the paleoproterozoic evolution of the southern nagssugtoqidian orogen and its southern foreland. key elements of this geodynamic evolution are also illustrated in fig. 101. the tonalitic to granodioritic intrusions in itillip ilua represent the very earliest dated paleoproterozoic event in the region. the c. 2500 ma intrusives clearly truncate the penetrative fabrics in the itilleq shear zone in the brittle domain, thus placing the formation of this shear zone into the (neo-)archaean era. available radiometric ages indicate that the itillip ilua – ikertooq area escaped major geological events for the next c. 450 my until the kangâmiut dyke swarm intruded the area from maniitsoq in the south to sisimiut in the north. as shown in fig. 6 (chapter 2), the available geochronological data clearly establish a kangâmiut dyke emplacement age of c. 2035 ma with a standard deviation of ±17 ma. given the regional extent of sampling, the total number of ages and the diverse techniques employed, this suite of dates confirms earlier arguments that this tectono-magmatic event marked the period of rifting that preceded nagssugtoqidian orogeny. however, the argument that the clusters of younger 40ar–39ar dates from kangâmiut dykes (1857 ± 14 ma and 1736 ± 20 ma; fig. 6) record slow cooling is not consistent with the overall metamorphic history of the orogen. extensive dating on country rock gneisses, schists and metasediments has established two metamorphic episodes for the nagssugtoqidian orogeny that cluster between c. 1850–1800 ma and c. 1780–1720 ma based on a range of techniques and isotopic systems (taylor & kalsbeek 1990; kalsbeek & nutman 1996; whitehouse et  al. 1998; nutman et al. 1999; connelly et al. 2000). these coincide, within their respective uncertainty envelopes, with the age data on metamorphic minerals from the kangâmiut dykes. a single sm–nd age based on the metamorphic minerals garnet, amphibole and plagioclase (1840 ma ± 80; stecher et al. 1998) further confirm the presence and extent of the older metamorphic event. based on these observations, we conclude that the cumulative data support an emplacement age of the kangâmiut dykes of 2035 ± 17 ma. these data are consistent with the interpretation that the dykes experienced and preserved evidence for the same younger metamorphic events as other rocks within the nagssugtoqidian orogen. the younger metamorphic dates can also be identified locally in the enclosing host rocks (fig. 5). it should be noted, however, that samples with nagssugtoqidian metamorphic ages are quite sparse south of itillip ilua, hence further refinement of the age and extent of the metamorphic overprints as recorded by host gneisses is not possible farther south. we envisage that the dykes were emplaced into an overall extensional setting or at least perpendicular to a minimum horizontal stress that could be overcome by the pressure of the ascending mafic magma bodies. previous workers interpreted the apparent conjugate set of shear zones (sinistral ne–sw and dextral wnw–ese zones) found locally in the host rocks and along the dykes as indicative of syn-tectonic dyke emplacement during overall nnw–sse compression. in this model, these structures guided the geometry of the intruding dykes. shears along dykes and in their marginal zones additionally suggest the presence of a nnw–sse-orientated horizontal stress field during or after dyke emplacement or both. it is important to note that these ‘conjugate shears’ are only documented between inussuttusup tunua and saqqap kangerluarsua and locally along kangerlussuaq. we favour an alternative model involving the possibly oblique opening of a rift basin into which the ikertooq supracrustals were deposited to take into account the 3–10% extension shown by the dyke swarm across the area (escher et al. 1975). we interpret the shears along the margins and in the interiors of dykes between inussuttusup tunua and saqqap kangerluarsua to be a distal response to the contractional deformation (nagssugtoqidian) responsible for the later deformation and metamorphism described in this study. this includes the high-grade deformation and metamorphism in the ikertooq area and formation of the ikertooq shear zone. the orientation of the extensional stress field responsible for formation of the proposed rift basin(s) and for the orientation of the kangâmiut dykes is poorly constrained at present. its original geometry has largely been obliterated by subsequent nagssugtoqidian orogenesis and more targeted work is clearly required to address this issue. available geochronological data indicate the presence of both archean and proterozoic sedimentary deposits in the nagssugtoqidian orogen (e.g. marker et al. 1999). the extent and volume of the supracrustal package in and along the ikertooq shear zone suggest that these sediments were likely laid down in a regional-scale, elongate basin (the ‘ikertooq basin’). available age data shows the youngest detrital grains to be c. 2100 ma (nutman et al. 1999) and the observation that relatively few dykes intrude these supracrustal rocks suggests that the dykes were mainly emplaced in the extended margins of the crustal blocks bordering the basin. additionally, it may also provide temporal constraints on the https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 87 of 96 geusbulletin.org age of deposition, thus suggesting that the main phase of deposition may even post-date dyke emplacement. more detailed discussion of the distribution, timing and significance of the supracrustal units in the nagssugtoqidian orogen is beyond the scope of the current paper. subsequent to emplacement, kangâmiut dykes between maniitsoq and itillip ilua suffered metamorphic overprints fig. 101 simplified cartoon illustrating the sequence of geological events described in the text for the study area between maniitsoq and ikertooq. a: stable archean crust intruded by tonalitic or granodioritic bodies at c. 2500 ma. these units cut existing shear zones (itsz) of unknown, but likely neoarchean age. iksz: ikertooq shear zone. itsz: itilleq shear zone. ausz: aujassoq–evighedsfjord shear zone. b: localised crustal extension, deposition of sediments in the ikertooq area and intrusion of the 2035 ± 17 ma kangâmiut dykes. c: large-scale south-directed thrusting affects basement gneisses, sediments and dykes and causes widespread metamorphism in the footwall as far south as maniitsoq. metamorphic ages recorded by the dykes are around 1850 ma. d: strong deformation and metamorphism in the area between itillip ilua and ikertooq with a northward-increasing gradient culminating in strongly sheared granulite facies rocks in maligiaq thrust over amphibolite facies rocks. most metamorphic ages are up to 100 myr younger than those obtained farther south towards maniitsoq. these ages may reflect a separate phase of tectonism or just the latest pulse in a protracted event starting with the thrusting in shown in panel c. pel: the erosional level presently exposed along the coastal transect discussed here. red box: approximate location of the red line shown in fig. 93. its relationship to the nagssugtoqidian front is not known nor is any implied. geographic locations refer to locations along the coastal transect. upper part of d, including the heavy black line, schematically represents observations farther east, closer to the greenland ice sheet (slightly higher crustal levels due to the gentle eastward crustal tilt). nagssugtoqidian front pel pel maligiaq ikertooq itillip ilua kangaamiut sermersuut maniitsoq itsziksz ausz itsz ausz itsz ausz itsz ausz a b c d archean quartzo-feldspathic gneisses, variably reworked during the paleoproterozoic shear zone (archean/proterozoic) supracrustal rocks in ikertooq (mainly paleoproterozoic) kangâmiut dykes present erosional level (along coastal transect) metamorphic isotherms (schematic) ~20 km ikertooq kangâmiut dyke swarm (2035 ± 17 ma) southerly thrusting at ~1850 ma ~2500 ma tonalitic/granodioritic intrusives truncating itilleq shear zone sisimiut north south fig. 3 ?? ?? https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 88 of 96 geusbulletin.org that range from thin amphibole rims on igneous clinopyroxenes to thoroughly recrystallised garnet-bearing metamorphic assemblages. the distribution of metamorphic overprints is very uneven, and all field observations and p–t calculations suggest that the isotherms and isobars of these overprints are gently undulating and largely sub-parallel with the present erosional surface (fig. 101d). the extent to which the metamorphic reactions approached completion may also have been controlled locally by the limited availability of reactants, such as fluids. different structural and metamorphic overprints were demonstrated in dykes and their host rocks between itillip ilua and maligiaq. in this area, the intensity of structural and metamorphic overprints increases gradually from south to north, culminating in granulite facies assemblages that were thrust southwards over amphibolite facies rocks in inner maligiaq. in contrast to farther south, the isotherms between itillip ilua and maligiaq are steeper and likely subparallel to the major thrust surfaces. it is notable that most dykes with preserved discordances show intrusion features, indicative of emplacement into brittle country rocks, which agrees with petrological considerations indicating that the present exposure level was at depths corresponding to ≤3 kbar during dyke emplacement (mayborn & lesher 2006). a mechanism by which these metamorphic overprints could have been generated has previously been proposed (mengel et al. 1995, 1996; mayborn & lesher 2006) and involves the wholesale loading and burial of the southern foreland area between ikertooq and maniitsoq by large-scale south-directed overthrusting. this is supported by the calculated metamorphic pressures from minerals in the overprinting assemblage of 7.5 kbar in the southern areas around maniitsoq and up to 10 kbar around ikertooq (fig. 100). available geochronological data from the dykes point toward two separate metamorphic events or pulses of a protracted period of foreland-directed thrusting. most ages of interpreted metamorphic origin are c. 1880– 1840 ma between maniitsoq and itillip ilua, whereas ages from north of itillip ilua to ikertooq are generally younger by 100–120 myr. a proposed explanation is that initial closure of the ikertooq basin and subsequent collision involved large-scale, southerly directed overthrusting, and that a later phase of more limited extent mainly affected the itillip ilua – ikertooq region, creating the ikertooq shear zone and the dramatic contact between granuliteand amphibolite-facies rocks in maligiaq. the younger age range notably correlates with structural and metamorphic events in the remainder of the nagssugtoqidian orogen. the observations summarised here are illustrated in fig. 101. the figure does not reflect the varying orientation of the kangâmiut dykes and only depicts schematically the difference in crustal levels exposed along the coast and towards the greenland ice sheet. similarly, the location of the nagssugtoqidian front (fig. 101d) is not known in the coastal transect focused on here and may well be represented by multiple zones of strong deformation in contrast to the geometry observed east of kangaamiut sermiat (e.g. fig 2). the metamorphism recorded between itillip ilua and maniitsoq (fig. 101c) is interpreted to represent an earlier event, separate to the one that generated the metamorphic and structural signatures in the itillip ilua – maligiaq area (fig. 101d). it should be noted that additional geochronological studies could very well show that these two events represent separate pulses of one protracted episode of large-scale south-directed thrusting. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 89 of 96 geusbulletin.org 10 concluding remarks in this contribution, we have described some of the key features of the structural, metamorphic and compositional aspects of the kangâmiut dykes between maniitsoq and ikertooq. the most pertinent of these include: 1. compositional homogeneity of the kangâmiut dykes across their mapped extent 2. geochemical similarities to basalts in tectonic rift environments 3. consistent intrusive ages around 2035 ma (where data are available) 4. uneven metamorphic overprints between maniitsoq and itillip ilua 5. a consistent structural and metamorphic gradient between itillip ilua and ikertooq 6. bimodal distribution of ages for metamorphic and structural overprints on the dykes the thermobarometric results from the dykes imply that north of the ikertooq shear zone, well preserved granulite-facies mineral assemblages developed at >800°c and about 10 kbar. south of the shear zone and inside the foreland towards itillip ilua, amphibolite facies mineral assemblages formed at approximately 700°c to 780°c and at pressures predominantly in the range of 6 kbar to 8 kbar. south of itillip ilua, data are sparser and display a slightly larger range than in the area between ikertooq and itillip ilua. geochronological data discussed in this contribution suggest that the metamorphic overprints preserved in the area around itillip ilua and to the south were generated during a thermotectonic event around 1850– 1800 ma. a subsequent event at 1780–1720 ma mainly affected the ikertooq – itillip ilua area, likely overprinting the earlier event. the former event affected a much larger area, however the significance of this must await more dense sampling and further study. lastly, we remain in awe of the work by geologists in the early phases of investigations of the precambrian in west greenland. their interpretations were based mainly on keen, but limited, field observations, however, the concepts they developed have largely stood the test of time. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 90 of 96 geusbulletin.org dedication in 2023, our dear friend, colleague and co-author william e. glassley suffered heart problems en route to denmark, where we were to spend a week planning a highly anticipated field season in greenland later in the summer. on arrival, bill was transferred to rigshospitalet in copenhagen where he passed away the following day on 19 march. bill’s involvement with greenland geology started in the late 1970s and never ceased: greenland became his destiny. he wrote about his fascination with greenland in the book “a wilder time – notes from a geologist at the edge of the greenland ice”, which was published in 2018. it is a uniquely poetic narrative in the great american ‘walden tradition’ of nature writing intertwined with science writing. another friend of bill’s, william griffin, wrote in a review of the book: “next time someone asks me why i am a geologist, i will just hand them this book”. we are lucky to have known bill and fortunate to have worked with him. he is sadly missed. acknowledgements jak received support from the danish natural science research council for field work in the itillip ilua – ikertooq region in 1974–1975 where he participated in the liverpool precambrian boundary programme led by juan watterson. jak and ks received support from the geological survey of denmark and greenland (geus) for field work in the itillip ilua – ikertooq region in 2016. fcm was supported in 1994–1998 by the danish lithosphere center nagssugtoqidian project funded by the danish national research foundation. weg received support from the united states national science foundation in 1980, aarhus university at various times in 2002–2015 and the carlsberg foundation in 2002. we wish to thank the reviewers for incisive and constructive comments and suggestions and acknowledge the geus bulletin editors for their patience and feedback. additional information competing interests the authors declare no competing interests. author contributions jk, fcm, weg: conceptualization, investigation, data curation, writing, review & editing. ks: structural model, review & editing. supplementary files two supplementary files are described in appendix a and b. these files are available at https://doi.org/10.22008/fk2/yo2mio. supplementary file s1: geochemical data for kangâmiut dykes. supplementary file s2: mineral chemistry from kangâmiut dykes. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ https://doi.org/10.22008/fk2/yo2mio korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 91 of 96 geusbulletin.org references allaart, j.h. 1982: geological map of greenland 1:500  000, sheet 2 frederikshåb isblink – søndre strømfjord. copenhagen: grønlands geologiske undersøgelse. allaart, j., & jensen, s. 1979: compilation of 1:500  000 reconnaissance mapping in the precambrian of the evighedsfjord – søndre strømfjord – itivdleq region, southern west greenland. rapport grønlands geologiske undersøgelse 95, 72–76. https://doi.org/10.34194/rapggu.v95.7651 armbruster, t. et al. 2006: recommended nomenclature of epidote-group minerals. european journal of mineralogy 18, 551–567. https://doi.org/10.1127/0935-1221/2006/0018-0551 árting, u.e. 2004: a petrological study of basic dykes and sills of assumed paleoproterozoic age in 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1–40. https://doi.org/10.34194/ geusm.v3.4596 van gool, j.a.m., connelly, j.n., marker, m. & mengel, f.c. 2002: the nagssugtoqidian orogen of west greenland: tectonic evolution and regional correlations from a west greenland perspective. canadian journal of earth science 39, 665–686. https://doi.org/10.1139/ e02-027 vernon, r.h. 2004: a practical guide to rock microstructure. 594 pp. cambridge university press. https://doi.org/10.1017/cbo9780511807206 vernon, r.h. & clarke, g.l. 2008: principles of metamorphic petrology. 446 pp. cambridge university press. waters, c.l., sims, k.w.w., perfit, m.r., blichert-toft, j. & blusztajn j. 2011: perspective on the genesis of e-morb from chemical and isotopic heterogeneity at 9-10°n east pacific rise. journal of petrology 52(3), 565–602. https://doi.org/10.1093/petrology/egq091 watterson, j. 1974: investigations on the nagssugtoqidian boundary in the holsteinsborg district, central west greenland. rapport grønlands geologiske undersøgelse, 65, 33–37. https://doi.org/10.34194/rapggu. v65.7383 whitehouse, m.j., kalsbeek, f. & nutman, a.p. 1998: crustal growth and crustal recycling in the nagssugtoqidian orogen of west greenland: constraints from radiogenic isotope systematics and u-pb zircon geochronology. precambrian research 91, 365–381. whitney, d.l. & evans, b.w. 2010: abbreviations for names of rock-forming minerals. american mineralogist 95. 185–187. https://doi. org/10.2138/am.2010.3371 willigers, b.j.a., mengel, f.c., bridgwater, d., wijbrans, j.r. & van gool, j.a.m. 1999: mafic dike swarms as absolute time markers in high-grade terranes: 40ar/39ar geochronological constraints on the kangâmiut dikes, west greenland. geology 27(9), 775–778. https://doi.org/10.11 30/0091-7613(1999)027%3c0775:mdsaat%3e2.3.co;2 winchester, j.a. & floyd, p.a. 1977: geochemical discrimination of different magma series and their differentiation products using immobile elements. chemical geology 20, 325–343. https://doi. org/10.1016/0009-2541(77)90057-2 windley, b.f. 1970: primary quartz ferro-dolerite/garnet amphibolite dykes in the sukkertoppen region of west greenland. in: newall, g. & rast, n. (eds): mechanism of igneous intrusion. geological journal special issue 2, 79–92. zeck, h.p. & kalsbeek, f. 1981: geochemistry of amphibolite facies metamorphism of a suite of basic dykes, precambrian basement, greenland. chemie der erde 40, 1–22. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ https://doi.org/10.1016/j.precamres.2021.106139 https://doi.org/10.1016/j.precamres.2021.106139 https://doi.org/10.1144/sjg18010049 https://doi.org/10.34194/geusm.v3.4596 https://doi.org/10.34194/geusm.v3.4596 https://doi.org/10.1139/e02-027 https://doi.org/10.1139/e02-027 https://doi.org/10.1017/cbo9780511807206 https://doi.org/10.1093/petrology/egq091 https://doi.org/10.34194/rapggu.v65.7383 https://doi.org/10.34194/rapggu.v65.7383 https://doi.org/10.2138/am.2010.3371 https://doi.org/10.2138/am.2010.3371 https://doi.org/10.1130/0091-7613(1999)027%3c0775:mdsaat%3e2.3.co;2 https://doi.org/10.1130/0091-7613(1999)027%3c0775:mdsaat%3e2.3.co;2 https://doi.org/10.1016/0009-2541(77)90057-2 https://doi.org/10.1016/0009-2541(77)90057-2 korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 95 of 96 geusbulletin.org appendix 1 geochemical data sources for the kangâmiut dykes supplementary file s1 (“geochemical data for kangâmiut dykes”) provides a compilation of geochemical data presented in this study. here we explain data sources as they are referred to in supplementary file s1 and as shown in fig. 46. beckmann (jack 1978): these data came from samples used in the paleomagnetic study by beckmann & mitchell (1976), and the full analyses were given in jack (1978). bridgwater (coll. 1972): these data represent samples collected by david bridgwater in 1972 and analysed at the chemistry facilities at the geological survey of greenland (now geus). some results based on parts of the data have been published previously (jack 1978; bridgwater et al. 1995). cadman, a.c., tarney, j., bridgwater, d., mengel, f., whitehouse, m.j. & windley, b.f. (2001): the petrogenesis of the kangâmiut dyke swarm, w. greenland. precambrian research 105, 183–203. https://doi.org/ 10.1016/s0301-9268(00)00111-x cadman et al. (2001), hansen (1989), jack (1978), korstgård (1980), mayborn (2000) and windley (1970): these data were all published previously. analytical details are given in the respective sources. korstgård et al. (this study): represents previously unpublished data from samples collected by j. korstgård and analysed subsequent to korstgård (1980; these data are referred to as ‘korstgård (unpubl.)’ in supplementary file s1), as well as samples collected by j. korstgård, w.e. glassley and k. sørensen in 2015 (referred to as ‘korstgård et al. (unpubl.)’ in supplementary file s1). unpublished data were analysed at bureau veritas laboratory services, vancouver, bc, canada. references beckmann, g.e.j. & mitchell, j.g. 1976: paleomagnetic and geochronological work in central west greenland. earth and planetary science letters 30, 269–280. https://doi.org/10.1016/0012-821x(76)90254-5 bridgwater, d., mengel, f., fryer, b., wagner, p. & hansen, s.c. 1995: early proterozoic mafic dykes in the north atlantic and baltic cratons: field setting and chemistry of distinctive dyke swarms. geological society, london, special publications 95(1), 193–210. https://doi.org/ doi:10.1144/gsl.sp.1995.095.01.12 cadman, a.c., tarney, j., bridgwater, d., mengel, f., whitehouse, m.j. & windley, b.f. 2001: the petrogenesis of the kangâmiut dyke swarm, w. greenland. precambrian research 105, 183–203. https://doi. org/10.1016/s0301-9268(00)00111-x hansen, b. 1989: en petrografisk og geokemisk undersøgelse af kangâmiutgangene, vestgrønland, 121 pp. msc thesis. geological institute, aarhus universitet, denmark. jack, s.m.b. 1978: the north atlantic proterozoic dyke swarm. 333 pp. phd thesis. liverpool university, uk. korstgård, j.a. 1980: metamorphism, deformation and chemical evolution of the kangâmiut dykes and their host gneisses across the ikertôq complex, the nagssugtoqidian mobile belt, west greenland. 236 pp. phd thesis. the university of liverpool, uk. mayborn, k.r. 2000: petrogenesis of the paleoproterozoic kangâmiut dike swarm, west greenland: implications for the tectonic history of northeast laurentia and the evolution of basaltic magmas, 318 pp. phd thesis. university of california, davis, usa. windley, b.f. 1970: primary quartz ferro-dolerite/garnet amphibolite dykes in the sukkertoppen region of west greenland. in: newall, g. & rast, n. (eds): mechanism of igneous intrusion. geological journal special issue 2, 79–92. https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ https://doi.org/10.1016/s0301-9268(00)00111-x https://doi.org/10.1016/s0301-9268(00)00111-x https://doi.org/10.1016/0012-821x(76)90254-5 https://doi.org/doi:10.1144/gsl.sp.1995.095.01.12 https://doi.org/doi:10.1144/gsl.sp.1995.095.01.12 https://doi.org/10.1016/s0301-9268(00)00111-x https://doi.org/10.1016/s0301-9268(00)00111-x korstgård et al. 2024: geus bulletin 58. 8312. https://doi.org/10.34194/geusb.v58.8312 96 of 96 geusbulletin.org appendix 2 a note on the mineral chemistry data from kangâmiut dykes supplementary file s2 (“mineral chemistry from kangâmiut dykes”) contians preliminary results based on some of the mineral analyses presented in mengel et al. (1994, 1995, 1996, 1997). all analyses were performed with the jeol 733 superprobe at the department of geosciences and natural resource management, university of copenhagen, denmark. mineral abbreviations from whitney & evans (2010). cations listed in the table were calculated on the basis of pyroxenes, 6 oxygens; feldspars, 8 oxygens; micas, 11 oxygens; amphiboles, 23 oxygens; garnet, 12 oxygens; all others, 1 oxygen. all fe is assumed to be fe+2. for amphiboles, calculation of cations and site occupancy for display and thermobarometry follows hawthorne et al. (2012) using a spreadsheet from locock (2014). amphibole and plagioclase analyses used in thermobarometric calculations are indicated with an “x” in column g (“p/t source”). to simplify the table, metamorphic garnet, clinopyroxene and orthopyroxene analyses used for geothermobarometry are not indicated. for these minerals, we generally used all or most analyses and worked with averages. references hawthorne, f.c., oberti, r., harlow, g.e., maresch, w.v., martin, r.f., schumacher, j.c. & welch, m.d. 2012: nomenclature of the amphibole supergroup. american mineralogist 97, 2031–2048. https://doi. org/10.2138/am.2012.4276 locock, a.j. 2014: an excel spreadsheet to classify chemical analyses of amphiboles following the ima 2012 recommendations. computers and geoscience 62, 1–11. https://doi.org/10.1016/j.cageo.2013.09.011 mengel, f., korstgård, j. & bridgwater, d. 1994: metamorphism of the kangâmiut dyke swarm, sw greenland: implications for development of the southern margin of the nagssugtoqidian orogen. igcp 275/371 meeting, nottingham, september 1994. terra nova 6, abstract supplement 2, 12. mengel, f., van gool, j. & marker, m. 1995: mafic dykes as monitors of orogenic development: an example from southern margin of the paleoproterozoic nagssugtoqidian orogen, west greenland. first dlc workshop on the nagssugtoqidian orogen in west greenland, proceedings, april 6–7 1995, copenhagen, 15–18. mengel, f., bridgwater, d. & hageskov, b. 1996: southern nagssugtoqidian foreland: tectonic and thermal evolution monitored by the proterozoic kangâmiut dyke swarm. dlc workshop on nagssugtoqidian geology – 1996, april 18–19 1996, copenhagen, 86–93. mengel, f., bridgwater, d. & hageskov, b. 1997: high metamorphic pressures in the orogenic margin and foreland – an example from the paleoproterozoic nagssugtoqidian orogen in west greenland. terra abstracts, terra nova 9, abstract supplement 1, 355. whitney, d.l. & evans, b.w. 2010: abbreviations for names of rock-forming minerals. american mineralogist 95. 185–187. https://doi. org/10.2138/am.2010.3371 https://doi.org/10.34194/geusb.v58.8312 http://www.geusbulletin.org/ https://doi.org/10.2138/am.2012.4276 https://doi.org/10.2138/am.2012.4276 https://doi.org/10.1016/j.cageo.2013.09.011 https://doi.org/10.2138/am.2010.3371 https://doi.org/10.2138/am.2010.3371 the kangâmiut dykes in west greenland: markers of the tectono-metamorphic evolution of the southern 1 introduction 2 regional geology 3 dykes and country rocks from maniitsoq to maligiaq 3.1 dykes and country rocks from maniitsoq to itillip ilua 3.1.1 isortoq to maniitsoq 3.1.2 maniitsoq to kangerlussuatsiaq 3.1.3 kangerlussuatsiaq to simiutaq 3.1.4 kangerluarsussuaq 3.1.5 inussuttusup tunua area 3.1.6 itillip ilua and surrounding areas 3.1.6.1 qaqqatoqaq and nuussuup qulaa 3.1.6.2 itilleq shear zone 3.1.6.3 dykes in the itilleq shear zone 3.2 dykes and country rocks of the ikertooq region (from saqqap kangerluarsua to maligiaq) 3.2.1 saqqap kangerluarsua (kangerluarsuk) 3.2.2 qeqertalik and avalleq 3.2.3 akulleq and maligiaq 4 igneous mineralogy and geochemistry of the kangâmiut dykes 4.1 igneous mineralogy 4.2 geochemistry of kangâmiut dykes in the sno and its southern foreland 5 metamorphic overprints on kangâmiut dykes from maniitsoq to itillip ilua 5.1 amphibole rims on ilmenite 5.2 amphibole rims on clinopyroxene 5.3 partial to complete replacement of clinopyroxene by amphibole 5.4 epidote growth in plagioclase 5.5 garnet growth on ilmenite and in amphibole rims on ilmenite 5.6 garnet growth in amphibole rims on clinopyroxene 5.7 titanite rims on ilmenite 5.8 disseminated garnet in dyke margins 5.9 distribution of metamorphic overprints from maniitsoq to itillip ilua 6 deformation and metamorphism of dykes and country rocks from saqqap kangerluarsua to maligiaq 6.1 microstructural changes 6.1.1 dykes metamorphosed under weak deformation 6.1.1.1 igneous microstructure 6.1.1.2 incipient granoblastesis, no deformation 6.1.1.3 intermediate granoblastesis, no deformation 6.1.1.4 advanced granoblastesis, no deformation 6.1.2 dykes metamorphosed under strong deformation 6.1.2.1 incipient granoblastesis, strong deformation 6.1.2.2 complete granoblastesis, strong deformation 6.1.3 conclusions regarding granoblastesis 6.2 metamorphism of dykes and country rocks 6.2.1 epidote zone 6.2.2 clinopyroxene-garnet zone 6.2.3 orthopyroxene zone 6.2.4 summary of metamorphic mineral assemblages between saqqap kangerluarsua and maligiaq 6.3 types of facies transitions in the itillip ilua – ikertooq region 6.3.1 facies transitions 6.3.2 metamorphic boundaries and magnetic anomalies 7 chemical changes in dykes and country rocks from nagssugtoqidian deformation and metamorphism 7.1 kangâmiut dykes 7.2 quartzo-feldspathic gneisses, amphibolites and supracrustal rocks 8 assessment of metamorphic conditions from observed mineral reactions and equilibria 8.1 geothermometry 8.1.1 methodology 8.1.2 results 8.2 geobaromometry 8.2.1 methodology 8.2.2 results 9 discussion 10 concluding remarks dedication additional information references appendix 1 geochemical data sources for the kangâmiut dykes references appendix 2 a note on the mineral chemistry data from kangâmiut dykes references figures fig. 1 simplified geological map of west greenland. solid black lines indicate the internal subdiv fig. 2 the coastal region between isortoq and sisimiut showing place names mentioned in text and key fig. 3 the classic sketch from escher et al. (1975) showing some of the pertinent features of the na fig. 4 distribution and orientation of kangâmiut dykes between kangaamiut and ikertooq. modified fro fig. 5 key published age data from the nagssugtoqidian orogen. a: 207pb-206pb age data measured on z fig. 5 key published age data from the nagssugtoqidian orogen. b: regional map showing the distribut fig. 6 all published dates for kangâmiut dykes, sorted by age. also shown are the computed means and fig. 7 field examples of dykes. a: archaean grey tonalitic to granodioritic polyphase layered gneiss fig. 8 field examples of dykes. a: 10 cm wide, discordant, undeformed kangâmiut dyke, hosted by arch fig. 9 a dense swarm of nne-trending kangâmiut dykes (yellow dashed lines) on the island itilliup qa fig. 10 field examples of dykes. a: coarse-grained garnet-bearing amphibolite from centre of c. 40 m fig. 11 field examples of dykes on itilliup qaqqaa, a small island on the south side of the mouth of fig. 12 satellite view of a 70 m wide e-w dyke, cut by a swarm of nne-trending kangâmiut dykes. at t fig. 13 field examples of dykes on simiutaq island, at the mouth of kangerlussuaq. a: 70 m wide e-w fig. 14 view ene from the north shore of kangerlussuaq towards sarfartoq (north wall of sarfartoq va fig. 15 satellite view of ne-trending dykes cutting an ese-trending dyke on the north coast of inuss fig. 16 orientation plots for dykes and host gneisses in areas along and adjacent to itillip ilua. d fig. 17 field examples of dykes from locations in and around itillip ilua. a: vein of grey tonalite fig. 18 field examples of dykes from saqqap kangerluarsua. a: veined grey gneiss on approaching a sh fig. 19 field examples of dykes from the head of itillip ilua. a: view towards east of the far weste fig. 20 examples of dyke geometries and relationships between dykes of different orientations in the fig. 21 rose diagrams showing orientation of kangâmiut dykes at itillip ilua. sinistral (orange) and fig. 22 penetratively deformed tonalitic to granodioritic amphibolite facies archaean gneiss. disrup fig. 23 field examples of dykes. a: steeply wnw-plunging lineation in the narrow part of a podded ka fig. 24 field examples of dykes from eqalugaarsuit on the north shore of itillip ilua a: sinistrally fig. 25 field examples of dykes from itillip ilua a: mylonitic zone in host gneiss adjacent to kangâ fig. 26 highly strained gneiss panels in the itilleq shear zone. a: detail of one high-strain panel fig. 27 large wnw-trending dyke cuts several smaller ene-trending dykes (fig. 20a). visible in the f fig. 28 features in podded kangâmiut dykes. a: train of several pods (yellow outline) adjacent to un fig. 29 sketch of podded dyke across a small peninsula at nuussuup qulaa on the south shore of itill fig. 30 intrusive features in podded dykes from the north shore of itillip ilua. a: field sketch of fig. 31 interpretive sketch of the 3d geometry of podded dykes. based on field observations of mostl fig. 32 strongly deformed tonalitic-granodioritic host gneiss adjacent to narrow part of podded dyke fig. 33 view from the widest part of the pod towards the narrow part (outlined in yellow dashed line fig. 34 intrusion features in kangâmiut dykes in itilleq shear zone. a: strongly sheared, straight g fig. 35 simplified geological map of the itillip ilua – ikertooq region. the ikertooq shear zone is fig. 36 field examples of gneisses from the eastern south shore of saqqap kangerluarsua. a: folded q fig. 37 field examples of dykes from the north shore of inner saqqap kangerluarsua a: mostly conform fig. 38 geological map (location in fig. 2), with stereographic plots that show the poles of planar fig. 39 field examples of supracrustal sequences (supra xx). a: steeply nw-dipping supracrustal sequ fig. 40 field examples of banded gneiss from inner qeqertalik. a: conformable metadyke in banded gne fig. 41 field examples from innermost qeqertalik. a: nw-verging folded banded gneisses looking west. fig. 42 field examples from north-east shore of avalleq. a: banded gneiss with older mafic bodies co fig. 43 field examples from around maligiaq. a: steeply north-dipping and orange weathering supracru fig. 44 dyke-gneiss relationships in northern ikertooq. a: steeply n-dipping metadyke is broadly con fig. 45 field examples from inner maligiaq. a: discordant and little deformed metadyke. hammer is 40 fig. 46 simplified geological map of the study area, showing the distribution and sources of analys fig. 47 photomicrographs from kangâmiut dykes with no or only minor metamorphic overprints. a: chill fig. 48 compositions of igneous and metamorphic minerals in kangâmiut dykes. a: compositions of all fig. 49 modal mineralogy of kangâmiut dykes from maniitsoq to ikertooq (based on point counting of t fig. 50 leucocratic centres in larger, zoned dykes. a: felsic material in centre of a >20 m kangâmiu fig. 51 bulk-rock compositional variations and classification of kangâmiut dykes. a: feo(tot)/mgo vs fig. 52 compositional variations and fractionation trends of kangâmiut dykes. a: pearce-type diagram fig. 53 mgo vs. major oxides for the kangâmiut dykes (all wt%). the distinct trends seen in mgo vs. fig. 54 mgo (wt%) vs. cao/al2o3. data points outline a cluster showing decreasing mgo with decreasin fig. 55 chemical variations for zr (ppm) and mg# from margins to centres of selected kangâmiut dykes fig. 56 trace and rare earth element (ree) characteristics of kangâmiut dykes. data sources in fig. fig. 57 early metamorphic overprints of amphibole rims on ilmenite. a: chilled margin of 30 m wide k fig. 58 early metamorphic overprints: amphibole rims on ilmenite and clinopyroxene. a: green amphibo fig. 59 early metamorphic overprints: partial to complete replacement of clinopyroxene by amphibole. fig. 60 early metamorphic overprints of epidote growth in plagioclase. a: coarse-grained, central ma fig. 61 early metamorphic overprints of garnet growth on ilmenite. a: delicate garnet rims along ilm fig. 62 a 70 m wide, foliated, recrystallised and zoned ne-trending dyke. two ilmenite grains with c fig. 63 early metamorphic overprints of garnet growth in amphibole rims on clinopyroxene. a: centre fig. 64 early metamorphic overprints of titanite rims on ilmenite. a: garnet-bearing, leucocratic ma fig. 66 early metamorphic overprints of disseminated garnet in margins of dykes. a: fine-grained dyk fig. 67 schematic to illustrate the metamorphic reactions that take place at ilmenite-plagioclase (t fig. 68 simplified map of the mineral assemblages described in section 3.1 showing the distribution fig. 69 examples of porphyritic microstructure in fine-grained dyke. a, b: porphyritic microstructu fig. 70 examples of sub-ophitic microstructure in coarse-grained dyke. a, b: coarse-grained dyke wi fig. 72 preserved sub-ophitic microstructure in dyke that is partly converted into a metadyke. a: 1 fig. 72 preserved sub-ophitic microstructure in dyke that is partly converted into a metadyke. a: 1 fig. 73 a dyke in which most clinopyroxene crystals have been replaced by amphibole. a: 1 nicol. b: fig. 74 a fine-grained dyke where most igneous minerals have been replaced by the metamorphic minera fig. 75 a completely recrystallised dykes with a faint planar fabric. a: sample 1286. partly igneous fig. 76 a strongly deformed dyke where lineation is much stronger than foliation. a, c: sample 1350a fig. 77 a strongly deformed dyke. sections cut perpendicular to the foliation. a: 1 nicol. b: x nico fig. 78 a strongly deformed dyke. a: sample 1300, 1 nicol. b: sample 1300, x nicols. in b, garnets s fig. 79 a completely recrystallised dyke. a, c: sample 1141a, 1 nicol, perpendicular to lineation. b fig. 80 dyke completely converted into a metadyke, or amphibolite, (sample 1127). a, c: 1 nicol. b, fig. 81 simplified geological map of the itillip ilua – ikertooq region, summarising metamorphic gr fig. 82 variations in dyke and metadyke mineralogy across the itillip ilua – ikertooq region (locat fig. 83 variations in country rock mineralogy across the itillip ilua – ikertooq region (location i fig. 84 mineral parageneses across the ikertooq region from south (bottom of diagram) to north (top fig. 85 metadyke in low amphibolite facies. a, b: foliated granoblastic microstructure with a planar fig. 86 examples of xx. a, b: large plagioclase porphyroblast in amphibolite facies metadyke. c, d: fig. 87 upper amphibolite facies metadyke. a (1 nicol), b (x nicols): good foliation with amphibole fig. 88 metadyke from clinopyroxene-garnet zone. a: (1 nicol) garnet (grt; pink), clinopyroxene (cpx fig. 89 titanite (ttn; centre of section) being replaced by ilmenite (ilm). a: 1 nicol. b: x nicols. fig. 90 granulite facies metadyke showing orthopyroxene (opx; pink), clinopyroxene (cpx; light green fig. 91 granulite facies metadyke. a: coexisting orthopyroxene, clinopyroxene, amphibole, garnet and fig. 92 data from four representative dyke samples between saqqap kangerluarsua and maligiaq. acf di fig. 93 simplified map of the itillip ilua – ikertooq area, showing the distribution of relative pr fig. 94 magnetic anomaly map for the itillip ilua – ikertooq area (modified from korstgård et al. 2 fig. 95 total intensity magnetic field anomaly map across the southern nagssugtoqidian orogen. the fig. 96 variations in dyke chemistry from south (area 01, left) to north (area 11, right) for those fig. 97 mobility of available lil elements from margin to centre in individual kangâmiut dykes in ar fig. 98 variations in k2o and rb in rocks hosting kangâmiut dykes from areas 07 to 11. data from qua fig. 99 distribution of samples for which temperature calculations are available (see table 3). in t fig. 100 distribution of metamorphic pressures (in kbar) between maniitsoq and sisimiut. grey circle fig. 101 simplified cartoon illustrating the sequence of geological events described in the text for tables table 1 geochronological data from selected mafic dykes in west greenland table 2 geothermometers and geobarometers used in this study table 3 pressure-temperature (p-t) data used in this study. moedt et al. 2025: geus bulletin 59. 8391. https://doi.org/10.34194/gn7zyz56 1 of 7 data article | short a chemistry and microbiology data set for meltwater rivers in south-western greenland (2017–2021) sanne m. moedt1 , kristian k. kjeldsen1, anders r. johnsen2 , andreas p. ahlstrøm1 , christian n. albers*2 1department for glaciology and climate, geological survey of denmark and greenland (geus), copenhagen, denmark; 2department for geochemistry, geological survey of denmark and greenland (geus), copenhagen, denmark abstract meltwater rivers in greenland transport large quantities of freshwater from the greenland ice sheet and local glaciers to the ocean, significantly influencing marine ecosystems and global biogeochemical cycles. with accelerating ice melt due to climate change, understanding the biogeochemistry of these rivers is critical. here, we present a data set providing comprehensive biogeochemistry data from 28 meltwater rivers in south-western greenland. spanning a period from 2017 to 2021, it includes data on nutrients and other ions, trace metals, sediment, radio and water isotopes, microbiology and cyanotoxins, sampled during field campaigns in june and august–september. this data set offers valuable insights for research on glacial meltwater, biogeochemistry and microbiology, addressing key knowledge gaps in these fields. *correspondence: cal@geus.dk received: 03 feb 2025 revised: 28 mar 2025 accepted: 15 apr 2025 published: 29 may 2025 keywords: freshwater chemistry, meltwater, microbiology, trace metals, water chemistry abbreviations: cfu: colony forming units pes: polyethersulfone geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: catherine n jex (geus, dk) reviewed by: two independent anonymous reviewers funding: see page 7 competing interests: see page 7 additional files: see page 7 data collection the original purpose of the data collection was to explore water composition and quality, which influenced the selection of data and methods to focus on parameters relevant to understanding the chemical and physical properties of the water. this included analysis of sediment content, nutrient levels and trace metals, as well as the use of techniques suitable for assessing both the general composition and specific contaminants in the water. brief description of the geographical, geological and climate setting the 28 study sites, all located in south-western greenland, were accessed by boat to navigate the various fjord systems. the initial selection of the sites was based on remote sensing and gis data sets to mainly assess accessibility, tabular abstract geographical coverage south-western greenland (60.9–66.5 °n) temporal coverage 2017–2021 (field campaigns in june and august–september) subject(s) cryosphere, hydrology, soils and biogeochemistry data format(s) raw and analysed biogeochemical and geographical data as csv files. all data are available at https://doi.org/10.22008/fk2/d1y8uy sample collection and analysis water samples taken from well-mixed river sites (n = 28), which were analysed for physical, chemical and microbiological parameters. parameters sediment concentration, ions, trace metals, radio isotopes, water isotopes, counts of total heterotrophic bacteria, coliform bacteria and enterobacteriaceae, cyanotoxins and geographical parameters. related publications n/a potential application(s) for these data assessment of freshwater biogeochemistry from a range of meltwater rivers in south-western greenland. https://doi.org/10.34194/gn7zyz56 https://orcid.org/0000-0002-4814-9686 https://orcid.org/0000-0002-9245-475x https://orcid.org/0000-0001-8235-8070 https://orcid.org/0000-0001-7253-3509 mailto:cal@geus.dk https://creativecommons.org/licenses/by/4.0/deed.en https://doi.org/10.22008/fk2/d1y8uy moedt et al. 2025: geus bulletin 59. 8391. https://doi.org/10.34194/gn7zyz56 2 of 7 geusbulletin.org abundance and water quality, respectively, as outlined in the works of ahlstrøm et al. (2018) and kjeldsen et  al. (2019). this included catchment delineation (see supplementary fig. si-1 to si-3) and river slope generated using a 30 m digital elevation model, while total ice cover within each catchment was determined from a vectorised promice ice mask (citterio & ahlstrøm 2013), updated with sentinel-2 imagery from summer 2018, and the randolph glacier inventory version 6.0 (rgi6.0; pfeffer et al. 2014), which maps local glaciers and ice caps around the year 2000. for catchments covered by both data sets, the rgi6.0 estimate was used for local glaciers and ice caps, while the updated promice ice mask provided the ice-sheet extent. sample collection water samples were collected from a total of 28 meltwater rivers (fig. 1) during field campaigns in june and late august – early september (herein referred to only as september), carried out over a period spanning 2017–2021 (see supplementary file s1). one location, d34, was sampled twice because of doubts about the correctness of initial radioactivity analyses (2018) for this site. chemical analyses were therefore included for this site in 2021 while microbiological analyses were not. water samples were collected from well-mixed areas of the rivers, free from visible animal droppings, typically a few metres from the riverbank. the sampler wore single-use nitrile gloves disinfected with ethanol prior to fig. 1 map of south-western greenland with sample locations. blue sites include cyanotoxin data, while orange sites do not. black squares: towns/settlements. https://doi.org/10.34194/gn7zyz56 https://geusbulletin.org/ moedt et al. 2025: geus bulletin 59. 8391. https://doi.org/10.34194/gn7zyz56 3 of 7 geusbulletin.org collection. some analyses were conducted on the boat using a temporary laboratory setup, allowing for on-site processing of samples. samples for the analysis of anions, cations, metals, stable water isotopes and cyanobacterial toxins (microcystins and anatoxin-a fumarate) were initially collected in an acid-washed glass bottle and processed within 2 h after sampling. for anions, cations and alkalinity, subsamples were filtered using a 0.45 μm polyethersulfone (pes) filter (q-max, frisenette) and transferred to plastic vials that were stored cold until analysis within a few weeks. for trace metals (excluding hg), nitric acid was added to a 30 ml plastic bottle before sub-sampling to preserve samples. trace element samples were analysed in both unfiltered and filtered (q-max pes 0.45 μm) forms to capture potential sediment interactions. total mercury was measured only in 2021, for which unfiltered and filtered samples were collected in acid-washed, airtight glass vials at the site, with acid added in the laboratory within 1–2 weeks. for stable water isotopes, 1.5 ml subsamples were filtered into glass analysis vials with septa. for microcystins, including nodularin, sodium thiosulfate was added by default to the 100-ml glass bottle prior to use. for anatoxin-a-fumarate (2017 only), the samples were collected directly in the river in a 1 l glass bottle. water for microbial counts was collected in sterile 50 ml centrifuge tubes, opened 10–20 cm below the water surface to avoid the surface film. sample analysis trace metals, radio isotopes and cyanotoxins were analysed by a commercial laboratory (eurofins miljø a/s; see table 1 for a complete list). here, trace metals were analysed according to ds/en iso 17294m:2016 icp-ms. cyanotoxins were analysed according to iso 20179 lc-ms/ ms for microcystin-lw, microcystin-lr, microcystin-rr, microcystin-yr and nodularin, and lc-ms for anatoxin a fumarate. radio isotopes were analysed according to nf en iso 10704 alpha/beta proportion for total alpha and beta activity, and so 13168:2015 liquid scintillation for tritium activity. total indicative doses were then calculated based on these measurements. it is important to note that the detection limits for radioactivity parameters reported by the commercial laboratory can vary from day to day due to variation in background radiation. anions and cations were analysed by ion chromatography (metrohm 819 ic) using a metrosep a 150/4.0 column for anions. stable water isotopes (δ18o and δ2h) were analysed on a picarro isotopic water analyser. the sediment concentration was determined as an average from three or four 1 l samples, which were each filtered through a 0.45 μm pre-weighed filter. the filter was then dried and weighed again to determine the sediment concentration in mg/l. next, the sediment was scraped off the filters and analysed for grain-size distribution on a malvern particle size analyser, using a subsample of at least 30 mg dried sediment. when less than 30 mg dried material was available, the sediment from all replicates was pooled before analysis. when the sediment concentration was less than 10 mg/l, sediment size distribution could not be determined. most microbial analyses were carried out on board the boat using sterile pipettes, alcohol-wiped gloves and an alcohol-wiped work surface. plate counts of table 1 summary table providing an overview of the data set. typical detection limits are noted. sample type parameter unit detection limit ph conductivity µs/cm temperature °c alkalinity meq/l 0.01 ions fluoride (f-) mg/l 0.005–0.04 chloride (cl-) mg/l 0.05 nitrate (no3 -) mg/l 0.05 phosphate (po4 3-) mg/l 0.05 sulphate (so4 2-) mg/l 0.05 sodium (na+) mg/l 0.05 potassium (k+) mg/l 0.05 calcium (ca2+) mg/l 0.05 magnesium (mg2+) mg/l 0.05 trace metals aluminium (al) µg/l 0.2 antimony (sb) µg/l 0.2 arsenic (as) µg/l 0.03 barium (ba) µg/l 1 lead (pb) µg/l 0.025 boron (b) µg/l 1 cadmium (cd) µg/l 0.003 chromium (cr) µg/l 0.03 cobalt (co) µg/l 0.04 copper (cu) µg/l 0.03 nickel (ni) µg/l 0.03 selenium (se) µg/l 0.05 zinc (zn) µg/l 0.3 mercury (hg) µg/l 0.001 sediment total sediment mg/l sand (>63 µm) % coarse silt (16–63 µm) % fine silt (2–16 µm) % clay (<2 µm) % radio isotopes total indicative dose msv/year 0.1 total alpha activity bq/l 0.02–0.06 total beta activity bq/l 0.1–0.4 tritium activity bq/l 5–10 water isotopes δ18o ‰ δ2h ‰ microbiology heterotrophs at 22 °c cfu/ml heterotrophs at 36 °c cfu/ml coliform bacteria cfu/ml enterobacteriaceae cfu/ml cyanotoxins anatoxin a fumarate µg/l 0.5 microcystin lw µg/l 2 microcystin lr µg/l 0.5 microcystin rr µg/l 0.5 microcystin yr µg/l 0.5 nodularin µg/l 0.5 https://doi.org/10.34194/gn7zyz56 https://geusbulletin.org/ moedt et al. 2025: geus bulletin 59. 8391. https://doi.org/10.34194/gn7zyz56 4 of 7 geusbulletin.org heterotrophic, colony-forming units were carried out using petrifilms (3m aqua heterotrophic count plate) for undiluted and 10-fold diluted samples, where the dilutions were prepared in filter-sterilised (0.2 μm) river water from the same location. the petrifilms were incubated in pentaplicates at 22°c (68 h) and 36°c (44 h) in accordance with ds/en iso 6222 to the extent possible when using custom-made, mobile incubators powered by 12 v lead-acid batteries. the temperature in each incubator was continuously monitored using a data logger. the incubators were placed on the deck. the 22°c incubator maintained a consistent temperature within the iso standard range. the 36°c incubator generally complied but occasionally dropped below the range during heavy wind. plate counts of faecal indicators were carried out using petrifilm for coliform bacteria (3m coliform count plate) and for enterobacteriaceae (3m enterobacteriaceae count plate). the petrifilms were inoculated with undiluted sample and incubated in pentaplicate at 36°c for 24 h according to ds/en iso 9308-1. after incubation, the petrifilms were stored at approximately 5°c and counted within 12 h. for the 2019–2021 field campaigns, colonies from the coliform petrifilms and selected enterobacteriaceae petrifilms were restreaked on macconkey agar in the laboratory and incubated overnight at 37°c to confirm growth at this temperature, and at 44.5°c to test for the presence of thermotolerant (faecal) e. coli. the isolates were identified by partial forward and reverse sequencing of the 16s rrna gene (27f and 1492r universal primers (weisburg et al. 1991), macrogen europe, 600–900 bp consensus sequences). the closest type material sequence matches were identified by blastn searches (zhang et al. 2000). data description and main features our data set contains water biogeochemistry data from 28 meltwater rivers in south-western greenland. the different sample types are accessible as separate csv files (supplementary files s2–s8). the data set is accompanied by a csv file containing the different sampling sites and their features (supplementary file s1), a read_ me text file describing the different data parameters and a pdf file containing three maps with the sites and their catchment delineations and ice cover (s9). here, we present an overview of the content of the data set. table 1 provides an overview of the different sample types collected at each site, the parameters analysed for each sample type, the units in which the parameters are presented and the detection limit for the analytical methods used. sediments the data set includes sediment concentration data (mg/l, supplementary file s3) from the glacial meltwater rivers, which carry suspended solids produced by the powerful erosion of basal rock as glaciers move across the terrain (herman et al. 2021). the greenland ice sheet contributes approximately 8% of the global fluvial export of suspended sediments to the ocean (overeem et al. 2017). these sediments also influence the colour and characteristics of greenland’s rivers and lakes, ranging from clear and dark to milky grey or brown (burpee et al. 2018). figure 2 shows the total sediment concentration divided into four size fractions: clay (<2 μm), fine silt (2–16 μm), coarse silt (16–63 μm) and sand (>63 μm). meltwater at most sites had a total sediment concentration of less than 200 mg/l. however, two sites in june, d10 and d11, had a sediment concentration of 444 and 1372 mg/l, respectively. the large differences in sediment concentrations in june and september observed at some sites, underscore the fact that samples need to be taken across the season to estimate an average sediment transport of a specific meltwater river. ions and trace metals inorganic water constituents, including nutrients, salts and trace metals, were analysed to assess river water characteristics, with particular attention to heavy metals. southern greenland’s geology features areas with high metal concentrations (hawkings et al. 2021), and the relatively unknown geology beneath the greenland ice sheet (macgregor et al. 2024) adds uncertainty to the presence of these metals in meltwater rivers. our data set contains a wide range of inorganic parameters to provide a broad overview of geochemistry in the region. supplementary file s2 shows elevated aluminium, nickel and chromium levels in unfiltered samples, likely due to partial extraction of these elements from sediment particles caused by the nitric acid that was added for preservation. in accordance with this, filtering the samples before the addition of nitric acid significantly reduced these concentrations. figure 3 highlights correlations between sediment concentrations and aluminium and nickel in filtered and unfiltered samples, with aluminium showing consistent fractions extractable with dilute nitric acid, while for nickel this varies across locations. it should be emphasised, that the extraction was not exhaustive and therefore the extractable concentrations cannot be used to conclude on the total amount of trace metals in the sediment particles. radio isotopes some areas in greenland contain radioactive minerals (keulen et al. 2014). therefore, the following radioactivity parameters were included: total indicative dose, https://doi.org/10.34194/gn7zyz56 https://geusbulletin.org/ moedt et al. 2025: geus bulletin 59. 8391. https://doi.org/10.34194/gn7zyz56 5 of 7 geusbulletin.org tritium and total alphaand beta-activity (supplementary file s4). no radioactivity was detected in any of the samples, besides low concentrations in d11 and d34. when d34 was resampled in 2021, values were below the detection limit. as detection limits of radioactivity parameters provided by the commercial laboratory vary from day to day, it is challenging to interpret results near these thresholds. water isotopes water isotope data (δ¹⁸o and δ²h) are provided in supplementary file s5. isotope fractionation occurs during precipitation, where colder temperatures result in fewer heavy isotopes (iaea 2001). inland and higher-altitude precipitation also have lighter isotopes, as heavier atoms are preferentially removed during transport. this results in distinct isotope signatures where the fig. 3 concentration of (a) aluminium, and (b) nickel for june and september in unfiltered and filtered samples versus sediment concentration of unfiltered samples. notice the logarithmic scales. three values of nickel were below the detection limit (<0.03 µg/l) and for these we report the detection limit only. fig. 2 sediment concentration (mg/l) and composition (clay, fine silt, coarse silt and sand, available only when sediment content was >10 mg/l) in meltwater rivers in june and september. d38, d39 and d40 were not sampled in september. note that some low concentration samples may be difficult to distinguish from the background line. to improve readability of the figure, the upper limit of the y-axis was set to 200 mg/l. in june, sites d10 and d11 had a total sediment concentration of 444 and 1372 mg/l, respectively. grain-size composition was not measured in 2017, hence there are no composition data for sites d27 and d38 in june despite their high sediment concentrations. https://doi.org/10.34194/gn7zyz56 https://geusbulletin.org/ moedt et al. 2025: geus bulletin 59. 8391. https://doi.org/10.34194/gn7zyz56 6 of 7 geusbulletin.org greenland ice sheet has the lightest signature (most negative δ¹⁸o-values), and coastal glaciers, coastal snow and summer rain tend to heavier signatures, with summer rain having the least negative δ¹⁸o-values. these signatures can indicate changes in the relative contributions of the various water sources to the meltwater river at different times. microbiology heterotrophic bacteria, coliform bacteria and enterobacteriaceae were detected at varying levels across sites and sampling periods (fig. 4 and supplementary file s6). coliform bacteria and enterobacteriaceae are generally used as indicators of faecal contamination but belong to the broad enterobacterales group that includes both faecal indicators and non-faecal environmental species. on 3m coliform petrifilm, coliforms are identified by acid production, indicated by a faint pink halo and gas production. however, distinguishing gas bubbles and halos under field conditions proved challenging, and all colonies were consequently counted as coliforms. this assumption was validated by replating coliform colonies and selected enterobacteriaceae colonies from the 2019–2021 field campaigns on macconkey agar at 37°c, and by 16s rrna gene sequencing of the colonies (supplementary file s7). all but two colonies were growth-negative at 44.5°c showing that they were not faecal e. coli. the colonies were identified to the genus level and belonged with few exceptions to common environmental enterobacterales genera, predominantly yersinia (47%), rahnella (19%) and serratia (16%). one of the thermotolerant colonies was a true escherichia as expected, whereas the other was a hafnia, an enterobacterales genus that generally should not grow under e. coli selective conditions, but is known to sometimes grow at a maximum temperature of 44°c (greipsson & priest 1983), which may explain its thermotolerance. cyanotoxins supplementary file s8 contains cyanotoxin data from meltwater rivers visited during field campaigns in 2017 and 2018. cyanobacteria, the dominant photosynthesising bacteria in arctic freshwater ecosystems, can produce toxins such as microcystins under certain conditions. these data were collected to assess the potential presence of cyanotoxins, as many meltwater rivers are connected to proglacial lakes where cyanobacteria are common. while no cyanotoxins were detected, the detection limit (table fig. 4 bacterial counts as colony forming units (cfu) per ml at the meltwater river sites in june and september for (a) heterotrophs at 22°c, (b) heterotrophs at 36°c, (c) coliform bacteria and (d) enterobacteriaceae. the detection limit was 0.2 cfu/ml. notice the differences in y-axis scale. https://doi.org/10.34194/gn7zyz56 https://geusbulletin.org/ moedt et al. 2025: geus bulletin 59. 8391. https://doi.org/10.34194/gn7zyz56 7 of 7 geusbulletin.org 1) was higher than the concentrations (0.005–0.4 μg/l) reported in a previous study using a highly sensitive immunochemical method that does not differentiate between microcystin and nodularin (trout-haney et al. 2016). acknowledgements the authors thank kisser thorsøe, karina hansen, bent hasholt, danielle hallé, øyvind a. winton and tina bundgaard bech for their technical and analytical support and field assistance, and erik palo jacobsen and the crew on m/s sterna (arctic boat charter) for safe navigation and valuable assistance in the field. the authors also thank the two anonymous reviewers whose comments improved the manuscript. additional information funding statement this study was supported by funding from promice, funded by the geological survey of denmark and greenland (geus) and the danish ministry of climate, energy and utilities under the danish cooperation for environment in the arctic (dancea) and the greenland integrated observing system (gios) under the danish agency for higher education and science. fieldwork and collection of samples was supported through collaboration with the government of greenland (naalakkersuisut). kkk acknowledges support from the independent research fund denmark (grant id 10.46540/3103-00234b). author contributions smm: writing – original draft, data curation, visualisation. kkk: conceptualisation, methodology, investigation, supervision, writing – review & editing. arj: methodology, investigation, writing – review & editing. apa: conceptualisation, methodology, investigation, writing – review & editing. cna: conceptualisation, methodology, investigation, writing – review & editing. competing interests the authors declare no competing interests. additional files ten additional files, including all data (s1_sampling_sites.csv, s2_ ions_and_trace_metals.csv, s3_sediment.csv, s4_radio_isotopes.csv, s5_water_isotopes.csv, s6_microbial_counts.csv, s7_coliform_enterobact_16rrna_id.csv, s8_cyanotoxins.csv) a readme file (read_me.txt) and supplementary figures s1, 2, 3 (supplementary_figure_si_1_3) are available at: https://doi.org/10.22008/fk2/d1y8uy. references ahlstrøm, a.p. et al. 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https://doi.org/10.1038/s41561-021-00753-w https://doi.org/10.1038/s43017-021-00165-9 https://doi.org/10.1038/s43017-021-00165-9 https://doi.org/10.22008/gpub/32038 https://doi.org/10.22008/gpub/32038kjeldsen https://doi.org/10.22008/gpub/32637 https://doi.org/10.22008/gpub/32637 https://doi.org/10.1029/2023gl107357 https://doi.org/10.1038/ngeo3046 https://doi.org/10.1038/ngeo3046 https://doi.org/10.3189/2014jog13j176 https://doi.org/10.3390/toxins8090256 https://doi.org/10.1128/jb.173.2.697-703.1991 https://doi.org/10.1089/10665270050081478 a chemistry and microbiology data set for meltwater rivers in south-western greenland (2017–2021) data collection brief description of the geographical, geological and climate setting sample collection sample analysis data description and main features sediments ions and trace metals radio isotopes water isotopes microbiology cyanotoxins acknowledgements additional information funding statement author contributions competing interests additional files references table table 1 summary table providing an overview of the data set. typical detection limits are noted. figures fig. 1 map of south-western greenland with sample locations. blue sites include cyanotoxin data, while orange sites do not. black squares: towns/settlements. fig. 2 sediment concentration (mg/l) and composition (clay, fine silt, coarse silt and sand, available only when sediment content was >10 mg/l) in meltwater rivers in june and september. d38, d39 and d40 were not sampled in september. note that some low concentration samples may be difficult to distinguish from the background line. to improve readability of the figure, the upper limit of the y-axis was set to 200 mg/l. in june, sites d10 and d11 had a total sediment concentration of 444 and 1372 mg/l, respectively. grain-size composition was not measured in 2017, hence there are no composition data for sites d27 and d38 in june despite their high sediment concentrations. fig. 3 concentration of (a) aluminium, and (b) nickel for june and september in unfiltered and filtered samples versus sediment concentration of unfiltered samples. notice the logarithmic scales. three values of nickel were below the detection limit (<0.03 μg/l) and for these we report the detection limit only. fig. 4 bacterial counts as colony forming units (cfu) per ml at the meltwater river sites in june and september for (a) heterotrophs at 22°c, (b) heterotrophs at 36°c, (c) coliform bacteria and (d) enterobacteriaceae. the detection limit was 0.2 cfu/ml. notice the differences in y-axis scale. hovikoski et al. 2023: geus bulletin 55. 8349. https://doi.org/10.34194/geusb.v55.8349 1 of 20 research article upper jurassic – lower cretaceous of eastern wollaston forland, north-east greenland: a distal marine record of an evolving rift jussi hovikoski1,2 , jon r. ineson1 , mette olivarius3 , jørgen a. bojesen-koefoed4 , stefan piasecki1,5,6 , peter alsen3 1department for geophysics and sedimentary basins, geological survey of denmark and greenland (geus), copenhagen, denmark; 2geological survey of finland (gtk), espoo, finland; 3department for geo-energy and storage, geological survey of denmark and greenland (geus), copenhagen, denmark; 4department for mapping and mineral resources, geological survey of denmark and greenland (geus), copenhagen, denmark; 5globe institute, university of copenhagen, copenhagen, denmark; 6retired abstract two drill cores spanning the upper jurassic – lower cretaceous succession in wollaston forland, north-east greenland, offer an insight into mud accumulation in an evolving distal fault block. previous studies have revealed the presence of long-lasting black mudstone accumulation extending through the oxygen-restricted early rift and rift climax phases (bernbjerg and lindemans bugt formations). here, we present a detailed description of the sedimentary succession extending into the late syn-rift settings (palnatokes bjerg and stratumbjerg formations). the results indicate that the kimmeridgian – lower volgian early rift-phase was characterised by suspension settling and millimetre-scale event deposition in a tectonically affected, prodeltaic offshore setting. the event-related depositional processes are recorded by starved wave ripples, scour-and-fill structures, putative mud-floccule ripples and mud-dominated gravity-flow deposits. during the middle volgian – ryazanian rift climax, the depositional environment evolved into a narrow half-graben that was detached from the proximal depocentre flanking the deltaic coastline, itself dominated by coarse sediment. the correlative sedimentary facies in the detached half-graben are bioclastic and pyrite-rich black mudstones, which document suspension settling and gravityflow or mass-wasting deposition in sub-storm wave-base slope and basin-floor environments. black mudstone sedimentation ended abruptly in the late ryazanian when the accumulation of condensed, bioturbated deep marine marls coincided with broader oceanographic reorganisation concomitant with waning rift activity in the west. deposition of red bioclastic mudstones with a common gravity-flow component characterised the hauterivian, potentially representing final draping of the submerged fault block crest. the top of the cored succession is demarcated by dark grey bioturbated mudstones of barremian age, reflecting the onset of regionally continuous deep marine mud accumulation in thermally subsiding basins. although superficially monotonous, the mudstone-dominated succession reveals a highly dynamic depositional system that reflects changing sediment transport processes during almost a full rift cycle. *correspondence: jussi.hovikoski@gtk.fi received: 03 mar 2023 revised: 12 sep 2023 accepted: 26 sep 2023 published: 21 dec 2023 keywords: late jurassic, early cretaceous, bernbjerg formation, lindemans bugt formation, palnatokes bjerg formation, stratumbjerg formation, mudstone, greenland abbreviations api: american petroleum institute units bi: bioturbation index gr: gamma ray wegf: wave-enhanced gravity-flow geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: karen dybkjær (geus, denmark) reviewed by: guy plint (western university, canada) and paul smith (oxford university museum of natural history, uk) funding: see page 18 competing interests: see page 18 additional files: none 1. introduction although mudstones form most of the earth’s sedimentary record, they remain less understood than most other rock types (schieber et al. 2007). this is unfortunate because mudstones form important archives of earth’s climatic and oceanographic processes and crises and are economically valuable, hosting petroleum source rocks and hydrocarbons and forming seals in hydrocarbon, groundwater and carbon storage reservoirs (potter et al. 2005). traditionally, mud has been regarded as being deposited via suspension settling under low-energy conditions (potter et al. 1980). however, particularly since the early 2000s, flume-tank experiments and detailed observations from the modern sea floor, and from ancient successions, have revealed that mud accumulation can take place during much https://doi.org/10.34194/geusb.v55.8349 https://orcid.org/0000-0001-6330-8713 https://orcid.org/0000-0003-0017-3705 https://orcid.org/0000-0003-3853-7543 https://orcid.org/0000-0001-5647-2769 https://orcid.org/0000-0002-7846-859x https://orcid.org/0000-0001-6218-9054 http://jussi.hovikoski@gtk.fi https://creativecommons.org/licenses/by/4.0/deed.ast hovikoski et al. 2023: geus bulletin 55. 8349. https://doi.org/10.34194/geusb.v55.8349 2 of 20 geusbulletin.org more varied hydrodynamic conditions than previously thought (kineke et al. 1996; wright et al. 2001; bentley & nittrouer 2003; macquaker & bohacs 2007; schieber et al. 2007; ichaso & dalrymple 2009; schieber & southard 2009; schieber & yawar 2009; macquaker et  al. 2010; ghadeer & macquaker 2011; plint et al. 2012; plint 2014; yawar & schieber 2017). most notably, such studies have shown that mud can be transported as, and deposited from, bedload transport, which can generate subtle sedimentary structures (e.g. cross lamination, normally graded lamina sets, structureless ungraded mud laminae and convergent lamination) that remain overlooked in the sedimentary record. this can be particularly true in outcrop studies where mudstone successions are typically poorly exposed relative to coarser facies and tend to appear homogenous due to weathering. the upper jurassic – lower cretaceous depositional record of eastern wollaston forland, north-east greenland (fig. 1), provides an exceptional window into mudstone deposition under changing paleoenvironmental conditions. the area experienced a protracted rift phase that started in the middle jurassic, reached a climax during the late jurassic and waned progressively during jurassic–cretaceous boundary times and through the early cretaceous (surlyk 1978, 1990, 2003). as a result of rift evolution, the basin geometry changed from a regionally continuous suboxic shelf setting to a series of narrow half-grabens that were subsequently affected by thermal subsidence and transgression. whilst the fill of the proximal half-grabens was coarse-grained, being characterised by major conglomeratic submarine fandelta systems (surlyk 1978, 2003; henstra et al. 2016), the distal fault blocks were mud-filled, stagnated and 18°w 16°w 76°n 72°n 22°w 20°w ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ g re en la nd traill ø hochstetter forland store koldewey wollaston forland jameson land 100 km 26°w 22°w24°w28°w 26°w 24°w " 74°45'n 74°50'n 74°55'n 74°40ʹn 74°35ʹn 74°20ʹn 74°15ʹn 74°25ʹn 74°30ʹn 19°30'w 19°0'w 20°0ʹw20°30ʹw brorson halvø-1 albrecht bugt daneborg p. f. k .f . h .f . rødryggen-1 w o l l a s t o n f o r l a n d k u h n ø young sund clavering ø s abine ø falske bugt major dykes/sills faults plateau lavas ice lakes rivers quaternary stratumbjerg and fosdalen fms palnatokes bjerg fm lindemans bugt fm bernbjerg fm pelion, jakobsstigen and payer dal fms permian pre-caledonian basement 10 km palaeogene borehole d .f . cretaceous jurassic triassic permian fault inferred deep-seated fault ■ ■ fig. 1 location maps. left: geological overview map of east and north-east greenland showing the location of the study area (red box) and the dominant late jurassic – early cretaceous faults. right: simplified geological map of wollaston forland showing the positions of the rødryggen-1 and brorson halvø-1 boreholes. k.f.: kuhn fault. p.f.: permpas fault. h.f.: hühnerbjerg fault. d.f.: dombjerg fault. the position of the coastline was primarily dictated by the dombjerg fault during the late jurassic, whilst the kuppel and hühnerbjerg faults were the most important structures controlling the distal half-graben architecture. reproduced from bojesen-koefoed et al. (2023a, this volume, fig. 1). https://doi.org/10.34194/geusb.v55.8349 http://www.geusbulletin.org/ hovikoski et al. 2023: geus bulletin 55. 8349. https://doi.org/10.34194/geusb.v55.8349 3 of 20 geusbulletin.org detached from the coastal gravity-flow systems (pauly et al. 2013; hovikoski et al. 2023; bojesen-koefoed et al. 2023, this volume; olivarius et al. 2023, this volume). the subsequent rift waning and thermal subsidence were associated with improved ventilation and deposition of deep marine mudstones. all these tectonostratigraphic phases are recorded in the fully cored boreholes rødryggen-1 and brorson halvø-1, which were drilled in two contrasting locations within a distal half-graben (figs 1, 2). full core recovery has yielded pristine preservation of the kimmeridgian to lower barremian mudstone-dominated succession, allowing detailed documentation of the sedimentary facies of the dominant, organic-rich black mudstone succession (bernbjerg and lindemans bugt formations), as well as the gradation to upper ryazanian bioturbated mudstones (palnatokes bjerg and stratumbjerg formations). the sedimentary facies of the black mudstone interval were recently summarised in hovikoski et al. (2023) in which the wider implications of this subsurface record are developed with respect to the development of anoxia in the region during rifting. the aim of this complementary paper is to present a comprehensive description of the depositional facies and processes in this unexposed and previously unknown distal marine setting. 2. stratigraphy and depositional setting the cored section covered by the rødryggen-1 and brorson halvø-1 boreholes spans the kimmeridgian to lower barremian interval, which is divided into four formations: (1) the bernbjerg formation, (2) the lindemans bugt formation (the new storsletten member; alsen et al. 2023, this volume), (3) the palnatokes bjerg formation (the albrechts bugt member and the rødryggen member) and (4) the stratumbjerg formation (fig. 2). the age of the bernbjerg formation is late oxfordian to early volgian, and it comprises a thick black fig. 2 stratigraphic models of the wollaston forland – kuhn ø area. left: the conceptual stratigraphic scheme of surlyk (2003) based on outcrop study. right: the revised scheme after drilling of the rødryggen-1 and brorson halvø-1 boreholes and incorporating changes introduced by surlyk et al. (2021). the revised succession is much more complete and richer in marine mudstone (brown colour) than previously envisaged. east greenl. tect. strat. seq.: east greenland tectono-stratigraphic sequence. u: upper/upper. m: middle/middle. l: lower/lower. lithostratigraphic abbreviations are as follows: al: albrechts bugt member. b: bernbjerg formation. ba: bastians dal formation. f: falskebugt member. j: jakobsstigen formation. l: laugeites ravine member. li: lindemans bugt formation. li (s): lindemans bugt formation (storsletten member). mu: muslingebjerg formation. n: niesen member. pa: palnatokes bjerg formation (young sund member). pay: payer dal formation. pe: pelion formation. r: rødryggen member. ri: rigi member. str: stratumbjerg formation. ug: ugpik ravine member. black vertical lines indicate the schematic location and stratigraphic extent of the rødryggen (r-1) and brorson halvø-1 (bh-1) boreholes. reproduced from bojesen-koefoed et al. 2023b (this volume). chronostratigraphywollaston forland – kuhn ø east greenl. tect. strat. seq. r falpa li ri n b j pe ug pay mu ba onlaps crystalline basementonlaps upper permian w e s n s n 2.6 2.5 2.4 2.3 2.2 2.1 hauterivian valanginian ryazanian volgian berriasian tithonian kimmeridgian oxfordian callovian alluvial/delta plain – paralic, sand-dominated coal shallow marine sandstones shelf transition – sandstones, mudstones, heteroliths o�shore/basinal mudstones deep marine sandstones deep marine conglomerates calcareous sandy marine mudstones red marine mudstones hiatus/condensed bathonian bajocian aalenian c re ta ce ou s ju ra ss ic lo w er u pp er m id dl e u m l u l u m l u m l u m l u l u m l l chronostratigraphywollaston forland – kuhn ø r ? al str bh-1 r-1pa no data li li (s)ri n b j pe ug pay mu ba onlaps crystalline basementonlaps upper permian w e barremian hauterivian valanginian ryazanian volgian berriasian tithonian kimmeridgian oxfordian callovian bathonian bajocian aalenian c re ta ce ou s ju ra ss ic lo w er u pp er m id dl e u m u l l u m l u m l u m l u l u m l l ? https://doi.org/10.34194/geusb.v55.8349 http://www.geusbulletin.org/ hovikoski et al. 2023: geus bulletin 55. 8349. https://doi.org/10.34194/geusb.v55.8349 4 of 20 geusbulletin.org mudstone shale succession (500–600 m) that accumulated in a tectonically influenced shelf setting (sykes & surlyk 1976; surlyk et al. 2021). previous sedimentological studies of outcrop sections have reported the presence of a dysoxic to anoxic or euxinic shelf setting with occasional storm influence during oxfordian–volgian times (surlyk & clemmensen 1975, 1983). similarly, the previous core-based study pointed to a generally sub-oxic prodeltaic shelf environment (hovikoski et al. 2023). rifting intensified during the volgian, which dissected the basin into a series of narrow, 10–30 km wide, westward tilted, fjord-like half-grabens (vischer 1943; surlyk 1978). major conglomeratic submarine fan-delta systems (lindemans bugt formation, rigi member; surlyk 1978; henstra et al. 2016) developed in the most proximal fault blocks reaching a maximum thickness of 2 km. the coeval paleoenvironmental development in more distal, eastern fault blocks remained poorly understood due to the lack of outcrops, but is well-recorded in the cores described here, which indicate deep basinal sedimentation and isolation from the coarsegrained proximal systems (hovikoski et al. 2023). the mudstone-dominated succession that accumulated in this setting is referred to the new storsletten member of the lindemans bugt formation (alsen et al. 2023, this volume). the rift climax started to wane in the west during the late ryazanian, leading to the deposition of the palnatokes bjerg formation (surlyk 1978, 1984, 1990, 2003). the formation changes coincided with transgression and improved ventilation in the water column and oceanographic change with improved communication between the boreal and tethyan realms (pauly et al. 2013). sand-dominated gravity-flow deposits (palnatokes bjerg formation, young sund member) continued to accumulate in the proximal fault block, whereas fossiliferous mudstones (albrechts bugt and rødryggen members) accumulated in basinal areas and on submarine block crests (surlyk 1978; hovikoski et al. 2018). the rødryggen-1 and brorson halvø-1 cores penetrate both the latter, fine-grained members, and new biostratigraphic data indicate a late ryazanian to hauterivian age for these deposits (alsen et al. 2023, this volume). in the brorson halvø-1 core, the palnatokes bjerg formation is gradationally overlain by a thin interval of uppermost hauterivian – barremian, sub-storm and wave-base bioturbated mudstones of the stratumbjerg formation (bjerager et al. 2020). this formation crops out from traill ø in the south to store koldewey in the north and reaches its maximum thickness of 270 m in the brorson halvø area; the stratumbjerg formation was previously referred to as the ‘mid-cretaceous sandy shale sequence’ (nøhr-hansen 1993). 3. methods the rødryggen-1 and brorson halvø-1 cores were sedimentologically and ichnologically described at a scale of 1:100. the sedimentological description included descriptions of lithology, grain size (visual estimation) and trends in grain size, primary and secondary sedimentary structures, bedding contacts and the identification of important stratigraphic surfaces. diagenesis and authigenic minerals are described in olivarius et al. (2023, this volume) and are integrated in facies descriptions. ichnological data comprise description of ichnogenera, trace-fossil assemblage, cross-cutting relationships and bioturbation index (bi of taylor & goldring 1993). the bi provides a description of the degree to which original sedimentary fabric has been destroyed due to biogenic processes. this classification scheme allocates a numerical value ranging from 0 to 6 – the values correspond to the percentage of bioturbation (cf. taylor & goldring 1993). undisturbed or non-bioturbated sedimentary fabrics are classified as bi 0 (0% reworked), while pervasively bioturbated media (100% reworked) are classified as bi 6. intermediate levels of bioturbation are characterised using bi 1–5 and are defined as follows: bi 1, 1–4% reworked; bi 2, 5–30% reworked; bi 3, 31–60% reworked; bi 4, 61–90% reworked; bi 5, 91–99% reworked (taylor & goldring 1993). locally, a lack of lithological contrast hindered accurate estimation of the degree of bioturbation. the age of the sediments is based on the biostratigraphy of alsen et al. (2023, this volume). 4. depositional facies the two cored sections spanning the bernbjerg, lindemans bugt, palnatokes bjerg and stratumbjerg formations are divided into seven facies (f1–7). the facies are grouped and described here according to the lithostratigraphic unit(s) in which they occur (figs 3, 4), as follows: •  f1, f2, f3 and f7 bernbjerg and lindemans bugt formations • f4 and f5: palnatokes bjerg formation •  f6: stratumbjerg formation 4.1 bernbjerg and lindemans bugt formations the studied deposits of the bernbjerg and lindemans bugt formations are divided into four facies (f1, f2, f3 and f7) and five subfacies (f1a, f1b, f1c, f2a and f2b) based on their sedimentological and ichnological properties. 4.1.1 facies f1: mudstone three subfacies (f1a–c) are recognised in the mudstone facies (fig. 5a–g). https://doi.org/10.34194/geusb.v55.8349 http://www.geusbulletin.org/ hovikoski et al. 2023: geus bulletin 55. 8349. https://doi.org/10.34194/geusb.v55.8349 5 of 20 geusbulletin.org fig. 3 sedimentological log of the rødryggen-1 core. note that the ammonite zones are here indicated with chronozone terminology, whereas in the text, ammonite zones are used. modified from hovikoski et al. (2023). chron. strat.: chronostratigraphy. tect. evo.: tectonic evolution. gr: gamma ray. comm.: comminus zone. scit.-wheat.: sciitulus–wheatleyensis zones. groenlandic.: groenlandicus zone. 10 m b er nb je rg f m li nd em an s bu gt f m w ol la st on f or la nd g ro up h al l b re dn in g g ro up pa ln at ok es b je rg f m a lb re ch ts b ug t m b 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 210 220 230 clay silt vf f m c f m cvc sand pebble gr rødryggen-1 facies dt chron. strat tect. evo. ki m m er id gi an lo w er v ol gi an m id dl e v ol gi an up pe r r ya za ni an lo w er r ya za ni an up pe r v ol gi an up pe r v al an gi ni an lo w er v al an gi ni an g en tl e bl oc k ro ta ti on ri ft c lim ax la te s yn -r if t w an in g ri ft in w es t p r a p a/p r interpretation sediment starved, oxic basin centre in tilted fault block hemipelagic sedimentation dominated basin centre gravity �ow in�uenced basin centre wave-enhanced gravity �ow in�uenced o�shore zone of max. regression wave-enhanced gravity �ow in�uenced o�shore storm-wave in�uenced proximal o�shore f1c/f2 f1c/b f1b/f1c f1c/ f1b f1c /b f1a/b f1b f7/f1a f1a f1a f1a/f1b f1b/f2a f1b/f1a f1a/f1b f1a/b f1a/f2 f1c f2a bi 0-3 bi 0-4 bi 0-6 bi 0-4 bi 2-4 bi 3-6py py py py py an an an an an an py py py py py py py p bi 0-3 f1c f1c f1c f1c f1c f1c f1c f2? f2b f2a f2a f2b/f3 f3 f2/f3 f2a f2 f2 f2 f2 f2 f2 f3 f2/f1b f2/f1 f2 bi trace fossilsfossilssediment structures2200 eu do xu s zo ne a ut is si od or en si s zo ne el eg an s zo ne sc it .-w he at . h ud dl es to ne -p ec ti na tu s pr im us c om m . li os tr ac us -p se ud ap er tu m g ro en la nd ic .ex ot ic us pr im it iv us m ay nc iko ch i mfr concretion clay clast dip of strata50° bivalve wave ripple cross-lamination coal bioturbation intensity belemnite faults slump, contorted lamination fractures shell fragments plant fragments bi 0-6 ripple cross-lamination chondrites burrow mottling planolites thalassinoides zoophycos very �ne sandstone/coarse siltstone heterolithic interlamination calcite-cemented sandy mudstone ankerite and dolomite-cemented mudstone planar laminated mudstone mottled lamination pyrite ankerite ammonite onychites loading synaeresis crack scour-and-�ll (gutter cast?) helminthopsis nereites mud �occule ripple py an retrogradation aggradation progradation r a depositional trendsdt p ? phycosiphon incertum ?bi 0-3 siphonichnus f4 st or sl et te n m b (api) f1c f2/ https://doi.org/10.34194/geusb.v55.8349 http://www.geusbulletin.org/ hovikoski et al. 2023: geus bulletin 55. 8349. https://doi.org/10.34194/geusb.v55.8349 6 of 20 geusbulletin.org py py py bi 6 f7/f1 f7/f2 f7/f2 f7/f1 f1a/f2a f2/f1a f1 f2 f2 f2 f3 f1c f7 f2 f1 fs fs p p p p p r r p r p r r r p r fs? fs fs? fs? r r/a f1c/f2 f1c/f2 f1c/f2 f1c/f2 f1/f7 f1/f2a f1c/f2 f1c/f2 f2b f2a f2a f2a f2a f2a f2a f2a f2a f2 f2 f1c f3/f2b f3/f2b f2/f1 f1b/c f2 f3/f2b f1c f1c f1b bi 0? bi bi 0-2 bi 0-4 bi 0-6 bi 0-6 bi 6 bi-6 bi-0 zone of max. regression zone of max. �ooding bi 0-2 bi 0-2 ? bi 0-2 bi 0-2 bi 0? bi 0-3 bi 0-2 bi 0-5 bi 0-6 ? bi 0-4 bi 0-3 py py py py py py py py py py py py mfr mfr ? ? ? py py py py py py py py py py py py py py py py an an an an py show show py py dominated py py py py py py py py py py brorson halvø-1 15° 50° 50° 45° 20° 5° 5° 5° 5° 10° 5° 5° 5°–? 15° 90° 10°–? 10° 15° 5° 5–10° 5–10° 20–25° 20° 20° 15° 50° 80° 20° 15–20° 15–35° 15–20° m m -c m -s ca le s lu m ps dm -s ca le s lu m ps cm -s ca le s lu m ps (m ic ro fa ul ts ) (m ic ro fa ul ts ) dm -s ca le sl um ps 10–15° 10–20° 20–25° 20–25° 20–25° 0–5° 0–5° 0–5° 10–15° 5–10° 0–5° 0–5° 0–5° 0–5° 0–5° 0–5° 0–5° 5–10° ? 5–10° 110 120 130 140 150 160 170 180 190 200 210 220 225.7 0 m 10 20 30 40 50 60 70 80 90 100 clay silt pebble facies dt chron. strat. tect. evo.interpretationbifossils bedding plane dip u. r up pe r v al .. – m . h au t. major block rotation onset of major block rotation late synrift – waning rift block rotation minor block rotation minor block rotation m.v lo w er vo lg ia n ki m m er id gi an eu do xu s zo ne a ut is si od or en si s zo ne el eg an s zo ne w he at ly en si s zo ne (~ 1 cm) an fs fs fs fs bi 0-3 sub-wave base slope (tilted fault block) slope sub wave base deep sea o�shore o�shore sub-wave base slope oxidised subaqueous, deep sea high oxidised subaqueous, deep sea high sediment starved tilted fault block sub-wave base slope (mud turbidites) o�shore gravity �ow in�uenced o�shore prodeltaic o�shore prodeltaic o�shore prodeltaic o�shore prodeltaic o�shore prodeltaic proximal o�shore prodeltaic o�shore prodeltaic proximal o�shore silt-rich proximal o�shore proximal o�shore sub-wave base slope (mud turbidites) prodeltaic o�shore redeposited ooze o�shore o�shore? ? gr 0 250 b er nb je rg f m h al l b re dn in g g ro up w ol la st on f or la nd g ro up pa ln at ok es b je rg f m r ød ry gg en m b stratumbjerg fm ? u.h. u. v l. b ar re m ia n ? sand f4 f5a f5b f5a f5b f6 api li nd em an s b ug t f m a lb re ch ts b ug t m b fs sb/fs trace fossilssediment structures vf f m c f m cvc fig. 4 sedimentological log of the brorson halvø-1 core. modified from hovikoski et al. (2023). chronostratigraphic abbreviations: l.: lower. m.: middle. u.: upper. vol.: volgian. r.: ryazanian. v./val.: valanginian. h./haut: hauterivian. legend in fig. 3. https://doi.org/10.34194/geusb.v55.8349 http://www.geusbulletin.org/ hovikoski et al. 2023: geus bulletin 55. 8349. https://doi.org/10.34194/geusb.v55.8349 7 of 20 geusbulletin.org 4.1.1.1 subfacies f1a: massive to laminated clayey mudstone description: f1a is a rare subfacies type occurring mainly in the lindemans bugt formation (in the core intervals: rødryggen-1 c. 97–24.4 m, brorson halvø-1 41–38 m; figs 3, 4), where it typically forms centimetreto metrescale successions. it commonly occurs in association with f7 (slumps) and f1b (colour-banded mudstone), and less commonly with f2 (clay-silt heteroliths) in the bernbjerg formation (fig. 4, 178 m). f1a has the finest grain size of the described facies and typically shows the highest gamma-ray (gr) values of the succession (>180 api; american petroleum institute units). it consists of dark grey to black, apparently structureless or faintly laminated clayey mudstone (fig.  5a). pyrite is locally common. generally, the deposits appear to be unbioturbated or contain rare diminutive chondrites, zoophycos and indistinct mottling (bi 0–3; fig. 5e, f). the transitional occurrences (f1a/f2a) in the bernbjerg formation are more commonly intensively micro-bioturbated with millimetre to sub-millimetre scale indistinct traces that are typically visible as laminae disruptions (bi 0–5). the estimation of bi and recognition of trace fossils was locally hampered by lack of lithological contrast and b c d e f g zo bm bm ~3 mm bm ~5 mm ~5 mm ~2 mm ~2 mm a fig. 5 facies f1. a: laminated black mudstone with common inoceramid fragments. f1a, rødryggen-1, 26.6 m. b–d: examples of ‘colour-banded’ pyrite-rich mudstone of f1b. fig. 5c illustrates slump-folded ankeriteand pyrite-rich mudstone. b: rødryggen-1, c. 64 m. c: brorson halvø-1, c. 47.5 m. d: brorson halvø-1, 46 m. e, f: bioturbated examples of f1ab. e: rødryggen-1, ~ 29 m. f: rødryggen-1, ~ 25.5 m. panel f represents a gradational interval from lindemans bugt to palnatokes bjerg formation. g: laminated to bioturbated ankeriteand dolomite-cemented mudstone. subfacies f1c, rødryggen-1, 76.6 m. bm: burrow mottling. zo: zoophycos. https://doi.org/10.34194/geusb.v55.8349 http://www.geusbulletin.org/ hovikoski et al. 2023: geus bulletin 55. 8349. https://doi.org/10.34194/geusb.v55.8349 8 of 20 geusbulletin.org poorly preserved core (rubble). finally, f1a is fossil-rich, bearing ammonites, belemnites and bivalves. interpretation: the lithology and sedimentary structures (e.g. laminated mud), concentration of marine fossils and ichnological properties are compatible with a marine, sub-storm wave-base, oxygen-deficient basinal environment characterised by hemipelagic suspension settling (see hovikoski et al. 2023 for redox data). the close association with f7 (slumps) in the middle volgian of the brorson halvø-1 section suggests a drowned slope setting in a tilted fault block at this location. the bioturbated, intermediate occurrences in the bernbjerg formation are interpreted as dysoxic offshore sediments based on ichnological and sedimentological characteristics as well as the stratigraphic occurrence of the facies. 4.1.1.2 subfacies f1b: colour-banded mudstone description: subfacies 1b is common in the lindemans bugt formation in rødryggen-1 (core depth 90–30 m). in the brorson halvø-1 core, f1b is rare, forming a thin interval in the middle volgian (core depth 43 m). in addition, a single occurrence of f1b transitional to f1c was recorded at around 152 m in the bernbjerg formation. in both of the latter instances, the subfacies is also associated with slumps (f7). f1b alternates gradationally with f1c and f1a, being an intermediate facies type between these two subfacies. on the gr log, f1b shows intermediate values commonly ranging between 100 and 150 api. this facies consists of interlaminated clayey mudstone, pyrite and ankerite-rich mudstone (see olivarius et al. 2023, this volume), which leads to the characteristic colour banding of this facies (fig. 5b, d). moreover, lensoidal laminae and possible low-relief cross-lamination occur locally, but their conclusive documentation is hampered by soft-sedimentary deformation and concretionary structures (e.g. pyrite nodules) that obscure primary micro-facies. bioclasts, consisting mainly of marine elements (e.g. onychites (belemnoid hooks), ammonites and bivalves), are also common. pyrite is abundant. f1b typically appears to be unbioturbated (bi 0) and shows small-scale soft-sedimentary deformation structures (micro-slumps consisting of contorted lamination and micro-scale loading structures). the determination of bioturbation intensity is hampered locally, however, by lack of lithological contrast in black mudstone-dominated intervals. interpretation: abundant small-scale soft sedimentary deformation, structureless mud laminae up to a few mm thick and local lenticular laminae suggest that deposition was probably partly derived from muddy gravity flows rather than being solely the result of hemipelagic suspension fallout. similarly, the alternation between stratified and structureless micro-facies suggests alternating transport processes, potentially driven by changing turbulence under decelerating mud flows (plint 2014). the gravity flows were probably facilitated by the increasing depositional gradient due to late early volgian – early ryazanian fault block development (surlyk 1978, 2003). the localised ankeritic laminae probably reflect alteration of rare primary laminae rich in bioclasts (e.g. calcispheres; see section 4.1.1.3; olivarius et al. 2023, this volume). considering the sedimentological and ichnological properties described here, coupled with the stratigraphic occurrence of the facies, f1b is interpreted to represent the accumulation of oxygen-deficient, basinal to slope mudstone in a tilted fault-block setting (see section 5. discussion). 4.1.1.3 subfacies f1c: ankeriteand dolomite-rich mudstone description: f1c is a common facies type in both core sections in the kimmeridgian and lower volgian (rødryggen-1 230–70 m, brorson halvø-1 220–48 m; fig. 5g). it occurs most commonly as decimetre-scaled intervals. on the gr logs, f1c is readily identified by anomalously low gr peaks (<50 api) that sometimes occur within an interval of overall high gr readings (at the base of an upward-coarsening interval) or at the very top of coarsening-upward successions. in addition, f1c occurs in trendless successions. f1c consists of ankeriteand dolomite-cemented mudstone, which characteristically shows interlaminated mudstone and bioclast-rich mudstone; the bioclasts are replaced by pyrite and ankerite (see olivarius et al. 2023, this volume). poorly preserved, their typical circular cross-section is suggestive of calcispheres, although some vase-shaped cross-sections resemble calpionellids. locally, f1c shows soft sediment deformation. pyrite and apatite are locally common (e.g. 89.8 m in rødryggen-1; olivarius et al. 2023, this volume). bioturbation intensity ranges from unbioturbated to intensive burrow-mottling (bi 0-6; e.g. brorson halvø-1, 150 m; fig. 5g). burrow mottling contains indistinct c. 1 mm wide, horizontal to sub-horizontal trace fossils, which produce a mottled fabric. these are tentatively assigned to zoophycos, because of (1) locally visible chevron-shaped structures, which may point to the presence of spreite, and (2) despite their diminutive size, they can in places be followed laterally through the width of the core, suggesting the presence of a lobe rather than an individual burrow. estimation of bioturbation intensity in f1c is locally hampered by concretionary structures related to ankerite and dolomite formation. interpretation: the enrichment of bioclastic material, authigenic minerals (phosphate) and locally high https://doi.org/10.34194/geusb.v55.8349 http://www.geusbulletin.org/ hovikoski et al. 2023: geus bulletin 55. 8349. https://doi.org/10.34194/geusb.v55.8349 9 of 20 geusbulletin.org bioturbation intensity is best explained by reduced sedimentation rates (condensation). this interpretation is supported by the occasional stratigraphic occurrence of f1c at the very top or base of parasequences, which could point to a reduced sedimentation rate during either transgressive events, tectonic reorganisation or both. however, local soft sediment deformation features and the close association with slumps (f7) suggest that some of the f1c layers (e.g. brorson halvø-1, 100–80 m) may represent redeposited calcareous ooze that was originally deposited as a pelagic drape on the eastern sediment-starved, incipient block crest, following the onset of block rotation. 4.1.2 facies f2: interstratified claystone, siltstone and sandstone description: f2 is the dominant facies type in the kimmeridgian – lower volgian bernbjerg formation in the rødryggen-1 cored section (234.5 to c. 97 m; fig. 3), where it typically forms several metres thick aggradational (i.e. trendless) successions, and is also well-represented in this formation in the brorson halvø-1 core (225.7–45 m; fig. 4). it is a transitional facies type with f3 (see section 4.1.3) and can be gradationally or erosionally interbedded with f3. in addition, f2 is locally intercalated with f1c. on the gr log, f2 typically represents aggradational successions displaying relatively uniform gr values (c. 140–160 api). locally, intervals dominated by f2 show stacked funnel-shaped gr patterns, a few metres thick (e.g. rødryggen-1, 175–148 m; fig. 3). f2 is most commonly unbioturbated, but locally, sporadic diminutive burrow mottling is recorded (bi 0–2). f2 is subdivided into two subfacies: f2a, consisting of laminated mudstone (fig. 6a, b), and f2b, which comprises interlaminated, very fine-grained sandstone or coarse siltstone and claystone and is characterised by lenticularity and basal erosional contacts (figs 6c–h, 7a, b). 4.1.2.1 subfacies f2a: parallel-laminated clay and silt description: f2a is a common facies type in the bernbjerg formation. it forms millimetreto centimetre-thick intervals occurring intergradationally between f1 and f2b. together, these subfacies form aggradational (trendless) successions, or subtle, up to a few tens of metres thick, upward-coarsening successions (e.g. rødryggen-1, 175–148 m; brorson halvø-1, 140–120 m). f2a consists of fine-grained heterolithic interlamination, typically parallel-laminated siltstone and claystone (fig. 6a). lamina pinch-outs and erosional scours are rare or absent. thin (c. 1 mm), tabular, normally graded siltstone–claystone couplets are present locally. moreover, the transitional expressions (f2a–f2b gradation) show increasing silt content and lamina thickness, and the appearance of normally graded lamina sets a few millimetres thick, which may laterally grade into scourbased, normally graded siltstone–claystone lamina sets (f2b). the deposits often appear unbioturbated or bear indistinct burrow mottling; rare diminutive zoophycos, chondrites and ?phycosiphon occur in places (bi 0–5; fig. 6a, b). disseminated pyrite and particularly coalified wood fragments are common locally. f2a differs from f2b in that it lacks (1) signs of erosion, (2) sandstone, (3) clear normal grading and (4) laminae convergence, pinch-outs or lenticularity. interpretation: in f2a parallel lamination dominates, with no indications of erosion. as in f1a, the deposits are bioturbated locally with possible diminutive zoophycos, chondrites and indistinct sub-millimetre scale burrow mottling, which is in line with a dysoxic, generally low energy environment (e.g. martin 2004; boyer & droser 2011; schieber & wilson 2021; see also hovikoski et al. 2023 for redox data). the normally graded, tabular siltstone–claystone couplets that lack erosional scour are interpreted as the most distal expression of thin mud-dominated gravity flows (see section 4.1.2.2 subfacies f2b). the transitional variants showing graded lamina couplets that may grade laterally into erosionally based siltstone–claystone lamina sets suggest that this type of interlamination is partly related to weak traction deposition. the interpreted depositional mechanism is the gravity-flow component of wave-enhanced gravity-flow (wegf) currents as described by macquaker et al. (2010; see section 4.1.2.2 for discussion). these wave-initiated flows possibly extended some distance below storm wave base as slope-maintained gravity-flow currents. in summary, considering the sedimentological and ichnological characteristics as well as the stratigraphic position, subfacies f2a is interpreted to record a low-energy dysoxic offshore environment that periodically experienced dilute, muddy gravity-flow events. 4.1.2.2 subfacies f2b: cross-laminated – lenticularly laminated silt and clay description: f2b is a common subfacies type in the brorson halvø-1 core. it forms millimetreto decimetre-scale successions and is commonly interbedded with f2a or f3. it usually occurs as a dominant subfacies in the top part of coarsening-upward successions (see section 5. discussion). f2b forms erosionally based upward-fining lamina sets a few millimetre thick, which may grade laterally into f2a over the width of the core. it is a broad subfacies dominated by interlaminated claystone–siltstone and lenses of siltstone, characterised by lamina truncations, low-relief erosional scours, lamina pinchouts and down-lapping to top-lapping lamina contacts (fig. 6c–h). https://doi.org/10.34194/geusb.v55.8349 http://www.geusbulletin.org/ hovikoski et al. 2023: geus bulletin 55. 8349. https://doi.org/10.34194/geusb.v55.8349 10 of 20 geusbulletin.org bm bm bm bm ? zo bm bm ? ph ? zo sy ~x cm~3 mm ~4 mm ~5 mm ~5 mm ~4 mm ~4 mm ~2 mm ~2 mm a b c a b c d e f g h fig. 6 facies f2. a: laminated mudstone with local burrow mottling (bm). f2a, rødryggen-1, c. 190 m. b: laminated to bioturbated mudstone (f2). biogenic structures are interpreted to include diminutive zoophycos (?zo) and phycosiphon (?ph). brorson halvø-1, c. 186 m. c: laminated silt and clay of f2b, showing normally graded laminae sets (black triangle), lenticular laminae and erosional contacts (yellow arrow). top of photo shows an interval of f2a with burrow mottling. rødryggen-1, c. 231.5 m. d: thin section image illustrating a typical tripartite division of a normally graded laminae set: a basal, erosionally based, cross-laminated very fine-grained sand or coarse silt lamina (unit a), parallel laminated silt and clay (unit b) and a clay-rich mud layer (unit c). f2b, rødryggen-1, 206.8 m. e: burrow mottling, visible as laminae disruption, in heterolithic interlamination. sy: synaeresis crack. white triangle: normally graded laminae set. brorson halvø-1, 106 m. f–h: facies examples illustrating laminae terminations (white arrows) and convergent lamination in f2b, suggesting the presence of ripple cross-lamination in mudstone. the lowermost cross-lamination in h contains sand and thus represents a transient example to f3. f: rødryggen-1, 232 m. g: brorson halvø-1, c. 150.5 m. h: brorson halvø-1, 215 m. https://doi.org/10.34194/geusb.v55.8349 http://www.geusbulletin.org/ hovikoski et al. 2023: geus bulletin 55. 8349. https://doi.org/10.34194/geusb.v55.8349 11 of 20 geusbulletin.org d b c e f ~5 mm ?ne ?ne ?ne ?ne ~10 mm ~5 mm ~5 mm ~5 mm ~2 mm a fig. 7 facies f2b and f3. a: normally graded sand-mud laminae sets with structureless mud laminae. yellow arrow points to syn-sedimentary deformation including small-scale slump folding. rødryggen-1, 233.69 m. b: a sand-filled scour (f3) grading laterally into mud-on-mud contact. the scour is interbedded with laminated silt and clay of f2b. brorson halvø-1, 135 m. c: mud-rich heteroliths (f2b), interbedded bioturbated intervals. whitedashed line marks the base of a structureless mud lamina. ?ne: possible nereites. brorson halvø-1, 185.5 m. d: ripple cross-lamination showing laminae offshoots typical for wave ripples (f3). brorson halvø-1, 147 m. e: cross-lamination with bioturbated top (f3). brorson halvø-1, 185 m. f: a scour-and-fill structure with ripple cross-laminated top (f3). the structure is interpreted as a storm wave-modified gravity deposit. brorson halvø-1, 183.5 m. https://doi.org/10.34194/geusb.v55.8349 http://www.geusbulletin.org/ hovikoski et al. 2023: geus bulletin 55. 8349. https://doi.org/10.34194/geusb.v55.8349 12 of 20 geusbulletin.org two different types of erosionally based, normally graded lamina sets and beds are present. type (1) comprises a few millimetre-thick siltstone–claystone couplets that show locally a tripartite microstructure: a basal micro-scoured contact below a millimetre-scale coarse siltstone or very fine-grained sandstone lamina (unit a; fig. 6c, d). this basal lamina shows lateral thickness variability, pinch-outs and may contain inclined sub-millimetre thick mud drapes. the basal unit a is abruptly overlain by millimetre-thick parallel-laminated claystone and siltstone (unit b in fig. 6d), which further grades into a structureless clayey mud-drape (unit c in fig. 6d) that may contain pyrite. type (2) comprises tabular, sharp-based siltstone–claystone laminae/beds. these beds are clay-dominated and appear to be composed of two components: a basal, c. 0.5–2 mm thick lenticular silt layer, which may contain inclined mud drapes and grades upwards into a structureless millimetreto centimetre-scale clay bed. the basal contact in the type 2 beds lacks evidence of prominent erosion. in addition to the upward-fining lamina couplets and triplets, f2b shows subtle changes locally in lamina angle, lamina truncations and on-lapping or top-lapping lamina contacts, particularly where the facies is interbedded with cross-laminated sandstone (f3; fig. 6f–h). locally, silt–clay laminae sets show subtle changes in lateral thickness (fig. 6h). soft-sediment deformation is also very common and occurs as millimetresto centimetre-scale micro-slump units (contorted z-shaped lamina). synaeresis cracks are rarely observed. the deposits commonly appear unbioturbated, though exhibiting locally a low-density or diversity trace fossil fabric similar to that of subfacies f2a (fig. 6e). the differences include lowered and fluctuating bioturbation intensity (bi 0–3) and increasing burrow diameter of burrow mottling in f2b (from sub-millimetreto millimetre-scale). moreover, in a few cases, an assemblage comprising nereites, ?chondrites and phycosiphon was observed at the top or in between event laminae or beds, at the gradation to f3 (fig. 7c). interpretation: the deposits are interpreted to be mainly related to various gravity-flow and wave-modified gravity-flow processes, but an influence from more sustained current systems is also possible. the heterolithic interlamination showing subtle changes in lamina angle, lamina truncations and down-lapping to top-lapping lamina contacts (figs 6g, h) suggests the presence of compacted, low relief mud-floccule ripples (macquaker & bohacs 2007; schieber et al. 2007; schieber & southard 2009). the silt–clay interlamination with laminae pinch-outs and lenticularity (fig. 6h) is similar to that generated by yawar & schieber (2017) in a flume tank experiment. their results showed that such clay–silt interlamination can form under similar flow velocity as mud-floccule ripples, but with a lower sedimentation rate. the type 1 erosionally based graded lamina set with the locally visible tripartite microstructure (figs 6c, d) is very similar to the wave-enhanced gravity-flow deposits described by macquaker et al. (2010). they proposed a three-phase flow model to explain the formation of these structures. in phase 1, wave-induced turbulence and resuspension form an erosionally based sand or coarse silt lamina (unit a). in phase 2, the increasing sediment concentration in the wave boundary layer damps turbulence, a pressure gradient develops and gravity flow begins. shear mixing at the base of the laminar flow and mixing with sediment already in suspension results in the deposition of interlaminated silt and clay (unit b). in phase 3, the flow energy wanes, and the lutocline collapses (i.e. the suspended mud cloud), flow stops and the deposits grade into a clay-rich layer (unit c). alternatively, such a structure could be deposited from a muddy gravity-flow current below wave base, which would also lead to turbulence damping as the flow decelerates and mud concentration increases (cf. baas et al. 2011). the wave-enhanced gravity-flow interpretation is supported here by the fact that the deposits either grade into or are interbedded with, wave-rippled sandstone (f3) and gutter casts (see section 4.1.3). the type 2 normally graded, tabular siltstone–claystone beds that lack prominent erosional scour at the base are interpreted as muddy gravity-flow deposits related to an increasing sea-floor gradient during episodes of block rotation. this interpretation is supported by the presence of intercalated slump deposits (f7). however, given the close association with deposits showing wave-influenced sedimentary structures, it is also possible that wave energy had a role in the formation of these deposits, and that the lack of unit b in these beds could be due to the dominance of clay, rapid flow deceleration or both. the locally occurring nereites-bearing ichnofabric and increasing burrow diameter suggest that oxygen levels increased at times during the deposition of f2b to f3. fluctuating bioturbation intensity and soft-sediment deformation are also compatible with elevated depositional rates and common small-scale event deposition. the abundant coaly material and plant debris point to distal fluvial influence. this is corroborated by hydrogen index and c30 desmethyl sterane trends, which reveal an elevated input of terrigenous organic matter (bojesen-koefoed et al. 2023, this volume). additionally, considering the sedimentological and ichnological characteristics as well as the stratigraphic position of the facies above dysoxic offshore sediments (f2a), f2b is interpreted to reflect deposition in a wave-influenced and fluvially sourced and tectonically https://doi.org/10.34194/geusb.v55.8349 http://www.geusbulletin.org/ hovikoski et al. 2023: geus bulletin 55. 8349. https://doi.org/10.34194/geusb.v55.8349 13 of 20 geusbulletin.org active proximal offshore environment with a variable sea-floor gradient. 4.1.3 facies f3: cross-laminated sandstone description: f3 is a subordinate, but recurring facies type in both cores. it forms a few millimetres to a centimetre thick and is commonly interbedded with f2. it is typically present in the upper part of coarsening-upward successions (see 5. discussion). f3 consists of ripple cross-stratified siltstone and very fine-grained sandstone. the ripples are characterised by asymmetrical to symmetrical profiles, irregular lower-bounding surfaces and common foreset offshoots that persist across the trough between ripples and peak again on the next ripple (fig. 7d). locally, the cross-laminated facies show a laterally variable erosional lower boundary forming scour-and-fill structures that are a few centimetres thick (figs 7b, f). the basal contact may be overlain by parallel-laminated, cross-stratified siltstone or very fine-grained sandstone onlapping the laterally limited basal scour. at the top, the siltstone–sandstone unit may be overlain by clay-rich mud drapes up to 1 cm thick (figs 7b, f). on the gr log, f3 is not always clearly expressed (due to the thin layers) but shows variable values of c. 80–120 api. f3 is commonly unbioturbated. in a single example, probable nereites intensively burrow the overlying mud drape (bi 0–5). interpretation: irregular lower bounding surfaces, bundling of laminae, foreset offshoots and locally occurring symmetrical ripple profiles suggest wave influence in the formation of the ripples (e.g. reineck & singh 1986). the laterally limited, relatively deep scour-and-fill structures showing onlapping lamina contacts over the basal erosional surface are interpreted to be storm-generated gutter casts (myrow et al. 2002). 4.1.4 facies f7: slumps description: in the brorson halvø-1 core section, f7 is a common facies type in the lower volgian of the bernbjerg formation and a dominant facies in the middle volgian lindemans bugt formation. it forms decimetre to metre-scaled intervals and consists of slumped mudstone (f1) or heterolithic sediments (f2; fig. 5c). interpretation: f7 is interpreted to record downslope mass wastage due to episodes of increased sea-floor gradients under block rotation. 4.2 palnatokes bjerg formation the palnatokes bjerg formation in the two cored boreholes consists of two facies: f4 of the albrechts bugt member and f5 of the rødryggen member. the latter only occurs in the brorson halvø-1 borehole and is subdivided into two subfacies. 4.2.1 facies f4: bioturbated calcareous mudstone description: f4 is a broad facies type, which forms the albrechts bugt member interval. it is dominated by bioturbated fossiliferous mudstone, which has a variable amount of sand and calcite cement in the matrix. locally, the deposits consist of variably bioturbated interlaminated grey mudstone and light grey calcareous mudstone. near the top of the albrechts bugt member, the facies appears structureless and contains shell fragments. f4 is generally characterised by a high fossil content. buchia bivalves are especially common. lithological accessories include abundant pyrite. f4 is typically intensively bioturbated (bi 3–6) with a relatively low-diversity trace fossil assemblage dominated by zoophycos and chondrites of various sizes or cryptobioturbation (fig. 8a, b). chondrites cross-cut or reburrow other trace fossils. secondary trace fossils include palimpsest grazing structures as well as rare siphonichnus and thalassinoides (figs 8b, c). interpretation: the ubiquitous biogenic reworking in this facies precludes detailed process interpretation. however, given the depositional setting, grain size and the nature of the underlying succession, it is likely that hemipelagic settling and dilute gravity-flow processes were prevalent. the change from the unburrowed black mudstones of the lindemans bugt formation to the fully bioturbated calcareous mudstones of the albrechts bugt member points to an environmental shift from an anoxic basin to an oxic basin with limited input of clastic sediment. the very low clastic sediment input is interpreted to be due to the late syn-rift paleogeographic configuration (compartmentalised basin) and eustatic sea-level rise (surlyk 1978). in summary, f4 is interpreted to represent sediment-starved deposits in a transgressed tilted fault-block setting. 4.2.2 facies f5: red calcareous mudstone description: f5 forms the rødryggen member interval in the brorson halvø-1 borehole. it consists of bioturbated red mudstone (fig. 8d–g), which has a variable carbonate content in the matrix; earlier works have shown that the red colour is due to haematite (alsen 2006). f5 is typically fully bioturbated (bi 6) exhibiting zoophycos, chondrites, siphonichnus, ?nereites and burrow mottling. facies f5 is divided into 2 subfacies. the first, f5a, consists of intensively bioturbated (burrow mottled), structureless red mudstone (fig. 8f). sand content is variable. the second subfacies, f5b, comprises bioclast-rich red mudstone and is characterised by a lower and fluctuating bioturbation intensity (bi 0–6). it contains sharp-based, shell hash beds that are a few centimetre thick (fig. 8d). in addition, dispersed inoceramus https://doi.org/10.34194/geusb.v55.8349 http://www.geusbulletin.org/ hovikoski et al. 2023: geus bulletin 55. 8349. https://doi.org/10.34194/geusb.v55.8349 14 of 20 geusbulletin.org fragments with borings (fig. 8g) and millimetre-scale rounded mud-clasts are very common locally. interpretation: the strongly composite biogenic fabric and the trace fossil assemblage point to a condensed, relatively deep, marine environment. the sharp-based shell-hash beds and locally incomplete bioturbation intensity in f5b are indicative of event deposition, probably due to gravity-flow events on a slope. earlier fieldwork has demonstrated that the rødryggen member is restricted to fault-block crests (surlyk 1978). given the sedimentological and ichnological evidence discussed here, f5 is interpreted to record sedimentation on an oxygenated submarine high in a deep marine setting (see 5. discussion). 4.3 stratumbjerg formation in the two borehole sections presented here, the stratumbjerg formation is represented only in the uppermost levels of the brorson halvø-1 core. 4.3.1 facies f6: bioturbated mudstone description: the interval referred to f6 in the brorson halvø-1 core is poorly preserved, consisting mainly of rubble. it comprises fully bioturbated (bi 6) light grey to greenish mudstone (fig. 8g). the deposits contain chondrites of different sizes, diminutive zoophycos, palimpsest grazing structures (helminthopsis and ?nereites) and rare thalassinoides. chondrites commonly re-burrow other trace fossils. inoceramus fragments and pyrite are common. the mudstone has a variable carbonate content, and concretionary beds are present locally. interpretation: the fine-grained lithology, the fully bioturbated fabric, the present ichnogenera and the lack of other forms suggest sub-storm wave-base basinal environment. the trace fossil content is not essentially different from that of the underlying palnatokes bjerg formation. in outcrop, the stratumbjerg formation is characterised by a regionally extensive grey mudstone succession with occasional thin sandy turbidites (bjerager et al. 2020). thus, the f6 mudstone facies in the uppermost brorson halvø-1 core probably represents a transitional variant recording the initial change towards renewed clastic sediment deposition in the area. 5. discussion the two drill cores provide an insight into mud accumulation in a distal fault block through almost a full rift cycle. whereas most of the previous research in the area has concentrated on characterisation of the coarse-grained gravity-flow systems in the coast-attached proximal fault block (e.g. surlyk 1978; henstra et al. 2016), the coeval sediments in distal areas have remained poorly documented. notably, distal sediments corresponding to the rift-climax phase were practically unknown prior to this drilling program. the facies recorded in the cores are diverse, reflecting sedimentation in the following settings: (1) storm-affected, oxygen-depleted deltaic shelf (bernbjerg formation; fig. 9a); (2) basin floor and slope in a dysoxic–anoxic half-graben (lindemans bugt formation; fig. 9b) and (3) sediment-starved deep half-graben under oxic conditions (palnatokes bjerg formation). finally, the top brorson halvø-1 core extends into the barremian, which records the gradation to (4) regionally continuous basinal mud accumulation under waning rifting (stratumbjerg formation). recently, hovikoski et al. (2023) summarised the depositional and tectonostratigraphic evolution of the black mudstone succession and broad-scale changes in bottom oxygenation during the rifting. here, we revisit the depositional evolution and further discuss the main facies. 5.1 bernbjerg formation (kimmeridgian – lower volgian) the sedimentary facies indicate that the bernbjerg formation represents an essentially aggradational muddy shelf succession that was fed by fine-grained fluvial systems (fig. 9a; see also surlyk & clemmensen 1983). in addition to macroscopic observations of common coalified wood fragments and plant debris, the prodeltaic nature is well-documented by organic geochemistry data, which indicate a prominent input of terrigenous organic matter (bojesen-koefoed et al. 2023, this volume; hovikoski et al. 2023). the seaway was influenced by south-orientated axial currents (hovikoski et al. 2023), which allowed along-coast dispersal of the river-supplied sediments. the upper cretaceous dunvegan formation of canada represents an analogous scenario (plint 2014). water depth ranged from a sub-storm wave-base offshore setting (f1; mainly laminated mudstone) to proximal offshore (f2b, f3; cross-laminated heteroliths and sandstone) and incipient slope environments (f7; slumps) below and above storm wave-base. the slope environment started to evolve, especially during the early volgian when rifting accelerated. elemental redox data and the distribution of bioturbation indicate fluctuating redox conditions in a generally hypoxic setting (see hovikoski et al. 2023). facies alternations occur on several scales ranging from millimetreto centimetre-scale facies alternations to depositional or tectonic cycles up to several tens of metres thick. the formation consists of several subtle, irregular transgressive–regressive and regressive cycles c. 10–30 m thick that individually comprise apparently chaotic facies alternations and rare, poorly developed higher-frequency depositional successions some metres https://doi.org/10.34194/geusb.v55.8349 http://www.geusbulletin.org/ hovikoski et al. 2023: geus bulletin 55. 8349. https://doi.org/10.34194/geusb.v55.8349 15 of 20 geusbulletin.org thick. an ideal, major upward-coarsening succession consists of f1, f2a, f2b and f3, reflecting progradation from dominantly low-energy dysoxic offshore conditions to proximal offshore with a higher frequency of depositional events. this ideal coarsening-upward trend is complicated by syntectonic influence, which affected depositional rate, bottom gradient, water depth and sediment source areas. particularly from the lower volgian and on, contrasting depositional rates are recorded by the two cores: the p. elegans – p. wheatleyensis ammonite chronozone interval is c. 30 m thick in the rødryggen-1 core, whereas the corresponding interval in the ~5 mm~5 mm~5 mm ~5 mm ~ 10 mm ~1 mm ~ 5 mm zo th ch zo zo zo ch ch ch sbb ? si bo bo a b c d e g f h fig. 8 a–c: examples of bioturbated marl (f4) characteristic of the albrechts bugt member. panel a is dominated by zoophycos (zo), whereas b illustrates common chondrites (ch) that cross-cut and re-burrow palimpsest traces. sbb: – spreite-bearing burrow potentially representing rhizocorallium. c shows potential pyritised siphonichnus (?si). a and b from rødryggen-1, interval around 20.5 m; c from rødryggen-1, 10.5 m. d–g: facies examples of red bioclast-rich mudstones of the rødryggen member (f5). panels d and e show occasional shell beds and scattered shell fragments in variably burrow-mottled fabric (f5b). visible traces include zoophycos. f illustrates intensively bioturbated structureless red mudstone (f5a). panels d, e and f from brorson halvø-1, 20 m, 22 m and 28 m, respectively. panel g is a thin section micrograph illustrating bored (bo) inoceramid fragments in sandy mudstone. brorson halvø-1, 9.7 m. h: bioturbated grey mudstone (f6) from the basal part of the stratumbjerg formation. two size classes of chondrites reburrowing thalassinoides (th). brorson halvø-1, 6.5 m. https://doi.org/10.34194/geusb.v55.8349 http://www.geusbulletin.org/ hovikoski et al. 2023: geus bulletin 55. 8349. https://doi.org/10.34194/geusb.v55.8349 16 of 20 geusbulletin.org brorson halvø-1 core is >100 m thick (alsen et al. 2023, this volume). due to these factors, as well as increasing clay laminae thickness in higher energy facies (f3), the coarsening-upward successions are commonly poorly expressed on the gr log. fining-upward successions related to transgressive phases are best developed in the lower part of the bernbjerg formation (kimmeridgian). these successions characteristically contain f1c beds, interbedded with f1b or f2. on the gr log, the successions are only weakly developed but are identifiable as serrated, generally increasing gr trends, punctuated by recurrent beds with low gr values (e.g. brorson halvø-1, 202–195 m). the stacking patterns of the major cycles suggest a maximum flooding zone in the upper part of the kimmeridgian, around 178 m in the brorson halvø-1 core. this is followed by an overall progradational trend, the zone of maximum regression being situated at around 100 m, in the lower volgian p. elegans zone. 5.1.1 sedimentary event beds and laminae although the lithology of the bernbjerg formation is generally fine-grained, indications of erosion, laminae-scale event deposition and traction currents are widespread. event deposition is recorded as wave ripples, scour-andfill structures (gutter casts), convergent lamination pointing towards putative mud floccule ripples (schieber et al. 2007) and associated silt–clay interlamination with pinchouts and lenticular laminae (figs 6, 7; yawar & schieber 2017). furthermore, normally graded siltstone–claystone beds, interpreted as muddy gravity flows, and probable wave-enhanced gravity-flow deposits occur frequently (macquaker et al. 2010; plint 2014). although similar structures may result from pure gravity-flow processes kf df hf a b df hf f1c, f2–3 f1ab f7 r-1 bh-1 r-1 bh-1 fig. 9 tectonostratigraphic–depositional scenarios of the wollaston forland area during the late jurassic – early cretaceous (modified from hovikoski et al. in 2023). a: kimmeridgian fluvially-sourced shelf influenced by along-coast sediment dispersal. the arrow indicates the main sediment transport direction, and red dots indicate borehole positions. df: dombjerg fault. hf: hühnerbjerg fault. b: middle volgian – ryazanian half-graben. the model shows major uplift along the dombjerg fault and the development of the coarse-grained fan deltas of the rigi member (lindemans bugt formation) in the coastal fault block (surlyk 1978, 2003). the fan delta-related gravity flows were blocked by the uplifted kuppel fault (kf) crest, which defined the western margin of the permpas/hühnerbjerg block(s). the hiatuses present in the brorson halvø-1 core indicate coeval uplift pulses along the hühnerbjerg block crest. r-1: rødryggen-1 borehole. bh-1: brorson halvø-1 borehole. subfacies f1c and f1ab along with facies f2–3 and f7 are shown. https://doi.org/10.34194/geusb.v55.8349 http://www.geusbulletin.org/ hovikoski et al. 2023: geus bulletin 55. 8349. https://doi.org/10.34194/geusb.v55.8349 17 of 20 geusbulletin.org below wave-base, their close association with gutter casts and wave ripples points to the involvement of stormwave processes. such structures are also missing from the lindemans bugt formation, which is interpreted to represent a sub-storm wave-base basinal and slope environment (see section 5.2). clay-rich mud layers can be up to 1 cm thick at the top of normally graded event beds. their structureless and seemingly ungraded nature coupled with common soft sedimentary deformation features (loading, slump-folding) suggest high water content and probably a fluid mud component in sedimentation. many sand-bearing event beds such as gutter casts pinch-out rapidly and are laterally equivalent to mudon-mud contacts. the mud-rich ripples range from well-defined cross-lamination with sand interlaminae (fig. 6h) to subtle silt–clay laminae sets showing lateral thickness variations and convergent laminae (fig. 6f, g). such occurrences are commonly observed adjacent to small, laminae-confined slump intervals, which may suggest that an increased slope gradient enhanced the mud-ripple development. overall, the observed facies are similar to those reported from other storm-affected mesozoic mudstone successions, including the cretaceous dunvegan formation (plint 2014) and the mowry shale (macquaker et al. 2010; lazar et al. 2022) from the western interior seaway. other examples include the jurassic cleveland ironstone formation, the whitby mudstone formation (ghadeer & macquaker 2011) and some intervals of the kimmeridge clay formation, uk (e.g. wignall 1989; macquaker & gawthorpe 1993). 5.2 lindemans bugt formation (middle volgian – late ryazanian) in the brorson halvø core, the base of the lindemans bugt formation is marked by a major hiatus or stratigraphic condensation as indicated by biostratigraphic data (alsen et al. 2023, this volume). similarly, the upper boundary of the formation is demarcated by a hiatus spanning from the middle volgian to the late ryazanian. these unconformities testify to further intensified tectonic activity in the wollaston forland area during the middle volgian (surlyk 1978, 2003). this rift climax resulted in tilted faultblock development and basin segmentation, lasting until the early late ryazanian. the brorson halvø-1 borehole is situated near the elevated hanging-wall crest of the permpas fault block about 30 km east of the main fault zone (the dombjerg fault). during the middle volgian, major conglomeratic fan deltas developed east of the dombjerg fault (surlyk 1978; henstra et al. 2016). the eastern limit of the submarine fan delta deposits was controlled by the kuppel fault, which defined the western margin of the permpas block (figs 1 and 9b). due to basin segmentation and changes in sediment source areas, the permpas block became isolated from the main focus of deltaic sedimentation and received progressively less terrestrial clastic sediment during rifting. this is directly reflected in the sedimentary facies exhibited by the brorson halvø-1 core, which suggest transformation into a non-deltaic, drowned oxygen-restricted slope (f7: slump) sub-storm wave-base slope (f1a, b: structureless to laminated mudstone) during the middle volgian. the presence of the two major hiatuses resulted from reduced accommodation caused by uplift of the block crest. 5.3 palnatokes bjerg and stratumbjerg formations the change from the lindemans bugt formation to the albrechts bugt member (palnatokes bjerg formation) in the rødryggen-1 core is gradational, occurring within 1 m of this boundary—a major hiatus in the brorson halvø-1 section. above this boundary, the clastic sediment input decreased, and the environment became oxygenated as indicated by bioturbation and elemental redox proxies (hovikoski et al. 2023). the change from black mudstone deposition to a ventilated basin with calcareous sedimentation is recognised supraregionally as an oceanographic change accompanied by the appearance of nannofossils with tethyan influence (pauly et al. 2013). the interpretation of limited clastic input and low sedimentation rate is supported by the increased fossil content, increased bioturbation intensity, composite ichnofabrics (cross-cutting, re-burrowing) and the nature of the ichnofauna (zoophycosand chondrites-dominated, low diversity ichnofabric). moreover, the estimated bulk depositional rates are low (hauterivian depositional rate c. 5.6 m/myr; hovikoski et al. 2023). in comparison to coeval outcrop data from the proximal fault block (hovikoski et al. 2018), land-derived turbidites are absent, indicating continued detachment from coastal depositional systems in the studied fault block. in the brorson halvø-1 section, the albrechts bugt member grades into the rødryggen member at around 30 m depth, marked by a shift in the sedimentary facies to red bioclastic mudstones. sedimentary facies suggest a more common gravity-flow component than in the albrechts bugt member suggestive of an increasing slope gradient and a rift pulse or multiple rift pulses. piasecki et al. (2020) recently described coarse clastic sediments of the falske bugt member (palnatokes bjerg formation) near the falkebjerg ridge, some kilometres east of the brorson halvø-1 drill site (fig. 1), indicative of intensified rift activity in this region during the valanginian–barremian. barremian syn-rift conglomerates are https://doi.org/10.34194/geusb.v55.8349 http://www.geusbulletin.org/ hovikoski et al. 2023: geus bulletin 55. 8349. https://doi.org/10.34194/geusb.v55.8349 18 of 20 geusbulletin.org not known from western fault systems (e.g. dombjerg fault), potentially suggesting that the rift climax persisted longer in the east. unlike the deposits of the falske bugt member, the coeval gravity flows of the rødryggen member comprise bioclasts and mud-clasts pointing to a limited or absent extraformational sediment source. this observation is compatible with localised footwall uplift and a compartmentalised basin. the setting gradually returned to an oxygen-restricted, sub-storm wave-base, deep basinal environment during the late hauterivian (basal stratumbjerg formation, f6). the sediments record decreasing biogenic carbonate accumulation and the renewed increase in clastic sediment input. in outcrop, the basal boundary of the stratumbjerg is variably developed, being either conformable, as observed in the brorson halvø-1 core, or hiatal, eroding into the bernbjerg formation (bjerager et al. 2020). in wollaston forland, this contact records waning rift activity, increased thermal subsidence and disappearance of rift-basin morphology. 6. conclusions the rødryggen-1 and brorson halvø-1 drill cores offer an insight into marine mud accumulation in an evolving distal fault block. the full core recovery, almost pristine preservation of the facies, stratigraphic continuity and the well-established biostratigraphic framework make the cores one of the best stratigraphic–sedimentological data points of the jurassic–cretaceous boundary and kimmeridge clay equivalent in northern high latitude regions. in particular, the kimmeridgian – lower volgian hypoxic shelf setting allowed exceptional preservation of subtle sedimentary structures in very thinly bedded mudstones recording distinct depositional events and traction currents and contributing to a general understanding of the processes governing mud accumulation. the cores document that black mudstone accumulation extended through the kimmeridgian – early ryazanian (late) early rift and rift climax phases (bernbjerg and lindemans bugt formations). the facies suggest that the early rift hypoxic prodeltaic shelf was characterised by suspension settling, starved wave ripples, scour-and-fill structures, putative mud floccule ripples and mud-dominated gravity-flow deposits. during the rift climax phase, the depositional environment evolved into a narrow half-graben characterised by bioclastic and pyrite-rich black mudstones. these deposits document hemipelagic suspension settling and gravity-flow or mass-wasting deposition in sub-storm wavebase, dysoxic, anoxic to euxinic slope and basin-floor environments. the ryazanian–valanginian late syn-rift setting experienced a supra-regional oceanographic change and improved ventilation, which is reflected in the deposition of deep marine marls (albrechts bugt member, palnatokes bjerg formation). condensed, red bioclastic mudstones with a common gravity-flow component characterised the hauterivian, which probably recorded an eastward shift in fault activity and final blanketing of the submerged fault-block crest (rødryggen member, palnatokes bjerg formation). the top of the cored succession is marked by the appearance of dark grey bioturbated mudstones of barremian age, recording the onset of regionally continuous deep marine clastic mud accumulation in thermally subsiding basins. thus, although superficially monotonous, detailed facies analysis of the mudstone-dominated succession exhibited by the rødryggen-1 and brorson halvø-1 boreholes reveals a highly dynamic depositional system that reflects shifting marine processes under varied hydrodynamic conditions, at different water depths and varied levels of bottom oxygenation during almost a full rift cycle. acknowledgements jette halskov, stefan solberg and jacob l. bendtsen are thanked for professional graphic support. guy plint and paul smith are thanked for constructive and insightful reviews. additional information funding statement funding for drilling of the rødryggen-1 and brorson halvø-1 boreholes and studies of the cores was provided by a consortium of oil companies and the geological survey of denmark and greenland (geus). author contributions jh: wrote the paper in co-operation with other authors. jh and ji: sedimentology. mo: mineralogy and diagenesis. jbk: organic geochemistry. sp and pa: biostratigraphy. competing interests the authors declare no competing interests. additional files none provided. references alsen, p. 2006: the early cretaceous (late 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https://doi.org/10.1144/gsjgs.150.1.0141 https://doi.org/10.1144/gsjgs.146.2.0273 https://doi.org/10.1016/s0025-3227(01)00140-2 https://doi.org/10.1016/s0025-3227(01)00140-2 https://doi.org/10.1016/j.sedgeo.2017.09.001 https://doi.org/10.1016/j.sedgeo.2017.09.001 upper jurassic lower cretaceous of eastern wollaston forland, north-east greenland: a distal marin 1. introduction 2. stratigraphy and depositional setting 3. methods 4. depositional facies 4.1 bernbjerg and lindemans bugt formations 4.1.1 facies f1: mudstone 4.1.1.1 subfacies f1a: massive to laminated clayey 4.1.1.2 subfacies f1b: colour-banded mudstone 4.1.1.3 subfacies f1c: ankeriteand dolomite-rich mudstone 4.1.2 facies f2: interstratified claystone, siltstone and sandstone 4.1.2.1 subfacies f2a: parallel-laminated clay and silt 4.1.2.2 subfacies f2b: cross-laminated lenticularly laminated silt and clay 4.1.3 facies f3: cross-laminated sandstone 4.1.4 facies f7: slumps 4.2 palnatokes bjerg formation 4.2.1 facies f4: bioturbated calcareous mudstone 4.2.2 facies f5: red calcareous mudstone 4.3 stratumbjerg formation 4.3.1 facies f6: bioturbated mudstone 5. discussion 5.1 bernbjerg formation (kimmeridgian – lower volgian) 5.1.1 sedimentary event beds and laminae 5.2 lindemans bugt formation (middle volgian late ryazanian) 5.3 palnatokes bjerg and stratumbjerg formations 6. conclusions acknowledgements additional information references figures fig. 1 location maps. left: geological overview map of east and north-east greenland showing the loc fig. 2 stratigraphic models of the wollaston forland kuhn ø area. left: the conceptual stratigraph fig. 3 sedimentological log of the rødryggen-1 core. note that the ammonite zones are here indicated fig. 4 sedimentological log of the brorson halvø-1 core. modified from hovikoski et al. (2023). chro fig. 5 facies f1. a: laminated black mudstone with common inoceramid fragments. f1a, rødryggen-1, 26 fig. 6 facies f2. a: laminated mudstone with local burrow mottling (bm). f2a, rødryggen-1, c. 190 m. fig. 7 facies f2b and f3. a: normally graded sand-mud laminae sets with structureless mud laminae. y fig. 8 a-c: examples of bioturbated marl (f4) characteristic of the albrechts bugt member. panel a i fig. 9 tectonostratigraphic-depositional scenarios of the wollaston forland area during the late jur research article | short larsen et al. 2023: geus bulletin 53. 8330. https://doi.org/10.34194/geusb.v53.8330 1 of 7 x-ray fluorescence (xrf) fingerprinting of palaeogene deposits in denmark nicolaj k. larsen*1 , kristian b.r. kristensen1 , marie-louise siggaard-andersen1 , claus heilmann-clausen2  , kurt h. kjær1 1globe institute, university of copenhagen, denmark. 2department of geoscience, aarhus university, denmark *correspondence: nicl@sund.ku.dk received: 07 oct 2022 accepted: 06 feb 2023 published: 10 mar 2023 keywords: denmark, geochemical fingerprinting, palaeogene, xrf analysis abbreviations: edxrf: energy-dispersive x-ray fluorescence geus: geological survey of denmark and greenland nr.: nørre pc: principal component pca: principal component analysis petm: paleocene–eocene thermal maximum xrf: x-ray fluorescence geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: karen dybkjær (geus, denmark) reviewed by: noël vandenberghe (ku leuven, netherlands), nicolas thibault (university of copenhagen, denmark) funding: see page 6 competing interests: see page 6 additional files: see page 6 abstract in this study, we test if cost-efficient x-ray fluorescence (xrf) analyses can be used to fingerprint palaeogene clay and marl deposits in denmark. a total of 67 samples from key sites in denmark have been analysed. our preliminary results indicate that it is possible locally within 10–30 km to distinguish between most of the palaeogene units, but on a regional scale across denmark, the units are not unique, and this probably reflects variations in clay mineralogy, grain size and calcareous content. accordingly, we suggest that a comprehensive reference database is now needed if the full potential of the method is to be utilised, and this will ultimately result in more reliable geological models. introduction in denmark, the surficial deposits (<100 m) are of a large societal interest (farming, forestry, natural resources and geotechnical properties), and they also comprise a vital groundwater reservoir. the surficial deposits consist of quaternary sediments (c. 10–30 m) above palaeogene sediments in northern and eastern denmark and neogene sediments in south-western denmark (binzer & stockmarr 1994; fig. 1). accordingly, pre-quaternary deposits are often encountered in boreholes during large infrastructure projects and in connection with groundwater projects. knowledge about the surface deposits mainly derives from more than 260 000 boreholes drilled in denmark in the last c. 125 years, and the data are available in the open-access jupiter database hosted by the geological survey of denmark and greenland (geus; hansen & pjetursson 2011). the jupiter database records the lithology and inferred age, whereas additional information that can be used to further discriminate sedimentary units (e.g. biostratigraphy) is only available for a very limited number of the boreholes. another complicating matter when making geological models is that the surface deposits have often been disturbed by glaciotectonics (jakobsen 2003). the heterogeneous surface geology in denmark thus presents a serious challenge when trying to make a geological model based on borehole data. one way to produce more reliable geological models is to make additional analyses (e.g. biostratigraphy and clay mineralogy) to characterise and discriminate the deposits (heilmann-clausen et al. 1985; nielsen et al. 2015). these procedures have, however, never been implemented routinely for geological investigations in denmark because the analyses are very time-consuming. in this study, we test a new cost-effective method to characterise and fingerprint palaeogene deposits in denmark using an x-ray fluorescence (xrf) analyser for discrete samples. we have analysed palaeogene units (mainly https://doi.org/10.34194/geusb.v53.8330 http://creativecommons.org/licenses/by-nc/4.0/ https://orcid.org/0000-0003-4171-9882 https://orcid.org/0000-0001-8986-1920 https://orcid.org/0000-0002-5894-7139 mailto:nicl@sund.ku.dk https://creativecommons.org/licenses/by/4.0/ larsen et al. 2023: geus bulletin 53. 8330. https://doi.org/10.34194/geusb.v53.8330 2 of 7 www.geusbul let in.org marls and clays) from key localities (fig. 1). our results show that the method can be used to discriminate most palaeogene units on a local scale (10–30 km) but lack the ability to differentiate the units on a regional scale across denmark. palaeogene deposits in denmark denmark is situated in the eastern part of the intracratonic cenozoic north sea basin (ziegler 1990). in most of paleocene and eocene times, the basin was inundated by relatively deep marine water that gradually became shallower and even subaerially exposed several times during the early oligocene (śliwińska et al. 2012; king et  al. 2016). overall, the palaeogene succession comprises a series of lithologically distinct units composed of early paleocene chalk, middle and late paleocene and eocene fine-grained clays and marls as well as oligocene silty clays (fig. 2). pre-cambrian crystalline rocks illite & chlorite cambrian sandstones palaeozoic limestone, shales & sandstones triassic–lower cretaceous sediments & sedimentary rocks upper cretaceous chalk paleocene–oligocene clay, silt & sand smectite miocene–pliocene sand & clay kaolinite danian limestone & chert 15° 58° 56° 54° 10° norway sweden denmark germany poland fur hinge kysing lundsgård klint tåsinge æbelø hesselbjerg mogenstrup nørre vissing juelsminde treldenæs albæk hoved ølst fig. 1 map of pre-quaternary deposits in denmark (based on sorgenfrei & berthelsen 1954). sample sites are divided into four local areas: limfjorden (purple), east jylland (yellow), lillebælt (red) and fyn (blue). https://doi.org/10.34194/geusb.v53.8330 http://www.geusbulletin.org larsen et al. 2023: geus bulletin 53. 8330. https://doi.org/10.34194/geusb.v53.8330 3 of 7 www.geusbul let in.org during the late cretaceous and early paleocene (danian), a chalk regime with the deposition of thick coccolithic chalk, bryozoan limestone and calcisiltite prevailed (thomsen 1995). the oldest siliciclastic deposits in denmark are the middle paleocene (selandian) lellinge greensand (up to 30 m) and the light grey kerteminde marl (c. 12–136 m), which in its upper part holds up to 50% reworked cretaceous chalk (thomsen 1995). above this follows two units deposited under increasing water depth: the grey, partly silicified æbelø formation (c. 16–57 m) and the late paleocene non-calcareous, varicoloured, very fine-grained holmehus formation (c. 12–40 m; nielsen et al. 1986; heilmann-clausen 1995). in the latest paleocene, the grey østerrende clay (up to 6 m) was deposited (nielsen et al. 1986). the paleocene–eocene transition coincided with a phase of intense activity in the iceland mantle plume leading to major basaltic volcanism centred on the rift zone in the norwegian–greenland sea (e.g. king et  al. 2016; ziegler 1990). during the paleocene–eocene thermal maximum (petm), the laminated and in part organic-rich stolleklint clay (c. 15–24 m) was deposited søvind marl fm viborg fm linde clay branden clay vejle fjord fm lillebælt clay fm røsnæs clay fm ølst fm fur fm stolleklint clay østerrende clay holmehus fm æbelø fm kerteminde marl danien limestone lo w er p al eo ce ne u pp er p al eo ce ne lo w er e oc en e m id dl e e oc en e u pp er e oc en e lo w er o lig oc en e u pp er o lig oc en e 23.8 28.5 33.7 37.0 49.0 55.5 60.0 65.6 lellinge greensand brejning fm east jyllandlimfjorden (n) (ø) værum mb 20 cm below 0 (h) værum mb 20 cm below +19 (h) værum mb 100 cm below top (h) (ø) r1 ± glauconite (h) r3 (h) r2 (h) r5 ± tephra (h) r4 (h) r6 upper/lower white bed + sapropel (h) l1 (h) l2 upper/lower white bed + sapropel (h) l3 (h,ø) l4 (h,ø) l5 + concretion (h,ø) l6 dark/light bed (h) pyt mb (h) kysing mb (k) viborg clay (k,ø) grundfør mb (h,ø) lower viborg clay (h) upper viborg clay (h) (f) knudeklint mb 10 cm below –33 (f) knudeklint mb -33 (f) knudeklint mb, concretion –24 to –29 (f) silstrup mb +19 (f) silstrup mb 10 cm below +19 (f) knudeklint mb 10 cm below –17 (f) skive clay (he) (m) ± glauconite (m) a: albæk hoved f: fur h: hinge he: hesselbjerg j: juelsminde k: kysing strand l: lundsgård klint m: mogenstrup n: nr vissing t: treldenæs tå: tåsinge æ: æbelø ø: ølst lillebælt fyn pyt mb (j) l4 (t) l3 (a) l2 (a,t) r5 (a) r4 (a) (a) ± silicified (æ) (j) (l) (tå) (tå) (tå) haslund mb 1-3 (tå) værum mb (tå) fig. 2 lithostratigraphy of the palaeogene deposits in denmark (after schiøler et al. 2007) and distribution of the 67 analysed samples. stratigraphy and ages of the lithostratigraphic units are based on heilmann-clausen (1995), clemmensen & thomsen (2005) and rasmussen et al. (2010). note that knudeklint and silstrup members are part of the fur formation. værum and haslund members are part of the ølst formation. stolleklint clay is a subunit of the haslund member. skive clay is a local facies of the branden clay. https://doi.org/10.34194/geusb.v53.8330 http://www.geusbulletin.org larsen et al. 2023: geus bulletin 53. 8330. https://doi.org/10.34194/geusb.v53.8330 4 of 7 www.geusbul let in.org (heilmann-clausen 1995). sometime after the petm, the character of the volcanism in the rift zone changed from effusive, forming huge flood basalts in east greenland and the faroe islands, to highly explosive, phreatomagmatic eruptions, spreading basaltic ash as far as 2000 km away in the northern tethys (e.g. stokke et al. 2020). in denmark, more than 180 ash layers were deposited in probably outer neritic waters. the ash layers are a few mm to 20 cm thick and occur in the clayey ølst formation, all over the north sea and most of denmark. a remarkable exception is nw jylland where the ashes occur in the diatomaceous fur formation (pedersen et al. 2011). after the deposition of the ølst and fur formations, sea-level rise led to a high eustatic sea level and the opening of the english channel and a marine connection across eastern europe to the peri-tethys in western asia (king et al. 2013). the hemipelagic, bathyal early eocene, mainly red and calcareous røsnæs clay formation (c. 3–28 m), and the early to middle eocene, mainly non-calcareous lillebælt clay formation (c. 40–100 m), were deposited in the present danish land area. the two formations are each subdivided into 6 members based mainly on differences in calcareous content and colour (heilmann-clausen et al. 1985). the røsnæs clay formation and lower members l1–l4 of the lillebælt clay formation resemble the holmehus formation in being varicoloured, very fine-grained and condensed. above these members follows the slightly siltier grey-green upper members l5 and l6 and the light grey middle to late eocene søvind marl formation (up to 90 m; heilmann-clausen et al. 1985). the søvind marl formation represents the last hemipelagic eocene sedimentary unit (thomsen et al. 2012) and is unconformably overlain by oligocene deposits of variable age. in the central part of the danish basin, sedimentation of marl continued locally to the eocene–oligocene boundary, and here, the søvind marl formation is usually overlain by the earliest oligocene viborg formation composed of up to 85 m thick dark grey finely micaceous silty clay, coarsening upwards to sandy silt (śliwińska et al. 2012; thomsen et al. 2012). above this formation usually follows the grey-green, silty, finely micaceous skive clay (80–90 m). the dark brown to nearly black, glauconitic and micaceous brejning formation was deposited in a marine, sediment-starved environment following a major fall in relative sea level during latest oligocene (rasmussen et al. 2010). on and south of the ringkøbing–fyn high, there is a major hiatus with only the oldest part of the søvind marl present, overlain by the brejning formation or younger units (rasmussen et al. 2010). materials and methods samples have been collected in type and reference sections (heilmann-clausen et al. 1985; heilmann-clausen 1995) and other biostratigraphically controlled sections (figs 1, 2). four samples from mogenstrup and nørre (nr.) vissing (supplementary file, table s1) are tentatively assigned to brejning and vejle fjord formations. the samples are not equally distributed across denmark or within the units but are the only samples available for this preliminary test of the method. from some units, we collected several samples if they were heterogeneous or contained, for example, concretions or ash layers. a total of 67 samples were analysed. each analysis was repeated five times to estimate the analytical uncertainty. after retrieval, the samples were dried in an oven at 90ºc for 24 h. the dry samples were loosened in a mortar and sieved to 0–500 μm before they were poured into 40 mm xrf sample cups. the samples were analysed using a bruker s2 puma energy-dispersive x-ray fluorescence (edxrf) spectrometer. the spectrometer was equipped with a 50 w x-ray tube with an ag anode. for each analysis, three different tube settings were used: no filter, a 250 µm cu filter and a 500 µm al filter in front of the x-ray tube for a total time of 400 sec. before each repeated analysis, the samples were gently stirred in the cup and lightly compacted. the analyses and data evaluation were performed using the bruker spectra. elements software for standardless xrf analysis. the spectrometer was calibrated using a glass disc (flx-k04 from fluxana) that was also used for drift monitoring. elements with insufficient counts, including some of the trace elements, have been excluded from the data set. results the xrf results show that the analytical uncertainty is minimal, and that there are variations in the abundance of elements between the palaeogene units (fig. s1 and table s1). elements like silicon (si), potassium (k), aluminium (al) and iron (fe) are abundant in all units but because they vary little, and they cannot be used to discriminate the units. in contrast, some units are rather unique as they have a high content of sulphur (s), such as the ash-bearing ølst, fur and stolleklint formations. other units contain higher amounts of calcium (ca), such as the kerteminde and søvind marl formations as well as the calcareous parts of the røsnæs clay formation (samples r3, r5 and r6 in fig. s1). barium (ba) and manganese (mn) are unique for the lillebælt clay and røsnæs clay formations and less abundant in the other units. however, none of the elements is unique and cannot be used to fingerprint a unit alone. principal component analysis (pca) of the data set groups the samples according to the formations that the samples come from (fig. s2). the first principal component (pc) represents varying proportions of lime and clay having positive loading from elements associated with lime (ca and sr) and negative loading from elements associated with clay (al, k, rb and si), whilst the second pc represents some diagnostic elements with positive https://doi.org/10.34194/geusb.v53.8330 http://www.geusbulletin.org larsen et al. 2023: geus bulletin 53. 8330. https://doi.org/10.34194/geusb.v53.8330 5 of 7 www.geusbul let in.org loading from ba and ni that are diagnostic for the lower lillebælt and upper røsnæs divisions and negative loadings from s and si that in high concentrations are diagnostic for ash layers and the fur formation, respectively. however, we do not consider pca further as a tool for fingerprinting because the groups are not sufficiently distinct and because some very different units are grouped together by pca. thus, results from a pca are not very useful to classify a new sample. instead, we have plotted the ratios between element concentrations for some of the most diagnostic elements in a 2d scatter plot, which makes it possible to differentiate most of the units (fig. 3 and fig. s3). we found that the ratios of k/fe and al/si plotted against 0.05 0.1 0.15 0.2 0.25 0.3 0.35 0.4 0.45 al/si 0.22 0.24 0.26 0.28 0.3 0.32 0.34 0.36 0.38 0.4 k /f e 0.32 0.34 0.36 0.38 0.4 0.42 0.44 0.46 al/si 0.2 0.25 0.3 0.35 0.4 0.45 k /f e æbelø fm (f) stolleklint (ej) stolleklint (l) holmehus fm (f) holmehus fm (ej) holmehus fm (li) østerrende clay (f) l1 (ej) l2 (ej) l4 (ej) l5 (ej) l6 (ej) l4 (li) l2 (li) l5 (ej) l3 (ej) l3 (li) al/si 0.1 0.15 0.2 0.25 0.3 0.35 0.2 0.4 0.6 0.8 1 1.2 k /f e fur fm diatomite (l) fur fm ash (l) ølst fm ash (f) ølst fm clay (ej) fur fm concretion (l) c b 0.25 0.3 0.35 0.4 0.45 al/si 0.15 0.2 0.25 0.3 0.35 k /f e kerteminde marl (f) pyt mb (ej) pyt mb (li) kysing mb (ej) søvind marl fm a 0.25 0.3 0.35 0.4 0.45 al/si 0.15 0.2 0.25 0.3 0.35 0.4 0.45 0.5 k /f e vejle fjord fm (ej) vejle fjord fm (l) skive clay (l) viborg fm (ej) brejning fm (l) brejning fm (li) f e 0.25 0.3 0.35 0.4 0.45 al/si 0.15 0.2 0.25 0.3 0.35 k /f e r2 (ej) r4 (ej)r3 (ej) r6 (ej) r5 (ej)r1 (ej) r4 (li) r5 (li) r5 ash layer (ej) d fig. 3 xrf plot (ratios of al/si and k/fe) from various palaeogene deposits. a: kerteminde marl and søvind marl fm. b: æbelø fm, holmehus fm, østerrende clay and stolleklint clay. c: fur fm and ølst fm. d: røsnæs clay fm divided into members r1–r6. e: lillebælt clay fm divided into members l1–l6. f: viborg fm, skive clay, brejning fm and vejle fjord fm. ej: east jylland, li: lillebælt, l: limfjorden and f: fyn. https://doi.org/10.34194/geusb.v53.8330 http://www.geusbulletin.org larsen et al. 2023: geus bulletin 53. 8330. https://doi.org/10.34194/geusb.v53.8330 6 of 7 www.geusbul let in.org each other resulted in the best differentiation for palaeogene units overall. in the palaeogene, there are two marl units: the paleocene kerteminde marl formation and the eocene søvind marl formation (fig. 3a). it is possible to differentiate the two formations, and it is even possible to distinguish between the two members, pyt and kysing, in the søvind marl formation. the next units comprise æbelø and holmehus formations, and østerrende and stolleklint clays from the upper paleocene and lower eocene, and it is also possible to distinguish between these (fig.  3b). in the æbelø formation, the xrf data are scattered, and this probably reflects the heterogenous nature of the unit, which is composed of silicified and nonsilicified clay. the more homogenous stolleklint clay and holmehus formation also have a scattered distribution, but this is probably because the samples were collected in different regions of denmark. the xrf data from the contemporaneous eocene ash-bearing diatomite (fur fm) and clay (ølst fm) are scattered, but the two formations can be clearly differentiated (fig. 3c). the large scatter within the formations is probably due to the heterogenous nature of the units, which contain c. 180 ash layers as well as numerous barium carbonate and calcium carbonate concretions. the xrf data show that it is possible to distinguish between all members in the overlying røsnæs clay and lillebælt clay formations within the same area, both in east jylland and in the lillebælt area (fig. 3d–e). however, there are changes in the composition within the same member between the different regions. the youngest palaeogene formations comprise the viborg, vejle fjord and brejning formations and the informal skive clay. within the different regions, the formations can be differentiated although there are significant inter-formational variations (fig. 3f). discussion overall, our preliminary results suggest that it is possible to differentiate most of the palaeogene units in the danish area using xrf fingerprinting on a local scale (e.g. limfjorden area) in combination with traditional facies analysis, whereas it is more difficult on a regional scale across denmark. in the lithologically homogenous units (e.g. holmehus, røsnæs clay, lillebælt clay and søvind marl formations), there is very little variation in the xrf results locally, and the fingerprinting method is very promising. in contrast, there is more variability locally in some of the more heterogenous units (e.g. fur, ølst and æbelø formations), which contain ash layers and silicified or calcareous intervals that are in part modified by diagenetic processes (fig. 3). on a regional scale across denmark, we observe a variability within both the homogenous and less-homogenous units, which is greater than the differences between units. accordingly, none of the palaeogene units has a unique xrf signature across denmark. the regional xrf variability across denmark probably reflects lateral facies variations in clay mineralogy, grain-size composition and calcareous content. in the norwegian–danish basin, the clay mineralogy reflects the different composition of the source rock (nielsen et al. 2015). for example, the high smectite content in paleocene and early eocene units reflects weathering of volcanic material, whereas abundant amounts of illite indicate a source area composed of metamorphic rocks. sorting of the clay minerals also influences the composition where larger particles of kaolinite are more abundant close to the shore, whereas smectite often dominates the central part of the basin (nielsen et al. 2015). conclusions in this preliminary study, we have used xrf to analyse 67 samples of palaeogene clayey and marly lithostratigraphic units from key sites in denmark for their content of selected elements. we find that it is possible locally, within 10–30 km, to distinguish between most of the palaeogene units, although some of them show large variability because of their content of ash, concretions and silicified intervals. on a regional scale across denmark, the units are not unique, and this probably reflects basin-wide variations in grain size, clay mineralogy and calcareous content. by establishing a comprehensive reference database, the full potential of this cost-effective method could be realised, and this can ultimately be used to make better geological models. this is now planned by analysing more samples from repositories and future geotechnical boreholes. acknowledgements we thank rené lyng sylvestersen from museum salling for providing samples from the fur formation. we also thank the two reviewers, nicolas rudolph thibault and noël vandenberghe, and the journal editor karen dybkjær for constructive comments and feedback. funding statement none declared. author contributions nkl and khk conceptualised the study. nkl, kbrk, mlsa and chc collected or supplied the material. kbrk and mlsa analysed the samples. all authors contributed to the interpretation of the data and in writing the manuscript. competing interests the authors declare no competing interests. additional files one supplementary file is available at https://doi.org/10.22008/fk2/ fchydq containing table s1 and figs s1–s3. https://doi.org/10.34194/geusb.v53.8330 http://www.geusbulletin.org https://doi.org/10.22008/fk2/fchydq https://doi.org/10.22008/fk2/fchydq larsen et al. 2023: geus bulletin 53. 8330. https://doi.org/10.34194/geusb.v53.8330 7 of 7 www.geusbul let in.org references binzer, k. & stockmarr, j. 1994: geologisk kort over danmark. 1:500.000. prækvartær-overfladens højdeforhold. danmarks geologiske undersøgelse. kortserie 44, 4–8 + maps. clemmensen, a. & thomsen, e. 2005: palaeoenvironmental changes across the danian–selandian boundary in the north sea basin. palaeogeography, palaeoclimatology, palaeoecology 219, 351–394. https:// doi.org/10.1016/j.palaeo.2005.01.005 hansen, m. & pjetursson, b. 2011: free, online danish shallow geological data. geological survey of denmark and greenland bulletin 23, 53–56. https://doi.org/10.34194/geusb.v23.4842 heilmann-clausen, c., nielsen, o.b. & gersner, f. 1985: lithostratigraphy and depositional environments in the upper paleocene and eocene 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greenland bulletin 12 keywords lithostratigraphy, biostratigraphy, north sea basin, palaeogene, neogene. cover complex fabric created by multiple small-scale sand intrusions (light) into dark mudstones – such enigmatic fabrics are commonly associated with the sand-rich units of the rogaland group in the siri canyon area, offshore denmark. the illustrated section of core is about 10 cm across and is from the lower tyr member (lista formation) in the cecilie-1b well (2346.8 m). photograph: jakob lautrup. chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, geological institute, university of aarhus; minik rosing, geological museum, university of copenhagen; finn surlyk, department of geography and geology, university of copenhagen scientific editors of this volume: jon r. ineson and martin sønderholm editorial secretaries: jane holst and esben w. glendal referees: paul van veen (norway) and robert o’b. knox (uk) illustrations: stefan sølberg digital photographic work: benny m. schark graphic production: knud gr@phic consult, odense, denmark printers: schultz grafisk, albertslund, denmark manuscript received: 29 august 2005 final version approved: 8 september 2006 printed: 29 june 2007 issn 1604-8156 isbn 978-87-7871-196-0 geological survey of denmark and greenland bulletin the series geological survey of denmark and greenland bulletin replaces geology of denmark survey bulletin and geology of greenland survey bulletin. citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 12, 77 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk and geografforlaget a/s filosofgangen 24, 1., dk-5000 odense c, denmark phone: +45 63 44 16 83, fax: +45 63 44 16 97, e-mail: go@geografforlaget.dk or at www.geus.dk/publications/bull © de nationale geologiske undersøgelser for danmark og grønland (geus), 2007 for the full text of the geus copyright clause, please refer to www.geus.dk/publications/bull 3 contents abstract . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5 introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7 geological setting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9 previous work . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10 material and methods. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 12 offshore and onshore lithostratigraphic nomenclature. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16 chronostratigraphy and biostratigraphy. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16 paleocene. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 20 eocene . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 20 oligocene . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 21 miocene . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23 lithostratigraphy . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 24 rogaland group . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 24 våle formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25 bor member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 30 lista formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 32 vile member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 36 tyr member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 39 ve member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 40 idun member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 41 bue member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 44 rind member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 45 sele formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 46 kolga member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 50 fur formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 51 balder formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 54 stronsay group . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57 horda formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57 hefring member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 63 westray group . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 63 lark formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 63 dufa member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 69 freja member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 71 acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 73 references . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 73 4 5 abstract schiøler, p., andsbjerg, j., clausen, o.r., dam, g., dybkjær, k., hamberg, l., heilmann-clausen, c., johannessen, e.p., kristensen, l.e., prince, i. & rasmussen, j.a. 2007: lithostratigraphy of the palaeogene – lower neogene succession of the danish north sea. geological survey of denmark and greenland bulletin 12, 77 pp. + 5 plates. as a result of a lithological, sedimentological and biostratigraphic study of well sections from the danish sector of the north sea, including some recently drilled exploration wells on the ringkøbing– fyn high, the lithostratigraphic framework for the siliciclastic palaeogene to lower neogene sediments of the danish sector of the north sea is revised. the sediment package from the top of the chalk group to the base of the nordland group is subdivided into seven formations containing eleven new members. the existing våle, lista, sele, fur, balder, horda and lark formations of previously published lithostratigraphic schemes are adequate for a subdivision of the danish sector at formation level. bor is a new sandstone member of the våle formation. the lista formation is subdivided into three new mudstone members: vile, ve and bue, and three new sandstone members: tyr, idun and rind. kolga is a new sandstone member of the sele formation. hefring is a new sandstone member of the horda formation. freja and dufa are two new sandstone members of the lark formation. danish reference sections are established for the formations, and the descriptions of their lithology, biostratigraphy, age and palaeoenvironmental setting are updated. __________________________________________________________________________________________________________ authors’ addresses p.s.*, j.a., k.d. & l.e.k., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. * present address: gns science, 1 fairway drive, avalon, p.o. box 30368, lower hutt, new zealand. e-mail: p.schioler@gns.cri.nz o.r.c. & c.h.-c., department of earth sciences, university of aarhus, høegh-guldbergsgade 2, dk-8000 århus c, denmark. g.d. & l.h., dong energy, agern allé 24–26, dk-2970 hørsholm, denmark. i.p. & e.p.j., statoil norway, forusbeen 50, n-4035 stavanger, norway. j.a.r., geological museum, university of copenhagen, øster voldgade 5–7, dk-1350 copenhagen k, denmark. 6 fig. 50 fig. 56a fig. 49 fig. 56b fig. 61 fig. 58 saxo-1 wessel-1 tordenskjold-1 eg-1 diamant-1 bertel-1 mona-1 karl-1 w. lulu-3,-1 cleo-1 augusta-1 amalie-1 tabita-1 gulnare-1 gwen-2 iris-1 baron-2 nora-1 elin-1w-1 ravn-1 falk-1 u-1 e-8 tove-1 john-flanke-1 alma-1 emma-1 edna-1 roxanne-1 ugle-1 frida-1 l-1 francisca-1 cecilie-1 connie-1 elna-1 siri-2 siri-3 sofie-1 floki-1 sandra-1 nolde-1 nini-1 nini-2 nini-3 d-1 vanessa-1 ibenholt-1 ida-1 r-1 c-1 k-1 f-1 inez-1 s-1 siri-1 v-1 g-1 deep-adda-1 adda-2,-1bo-1 north-jens-1 lulu-1 lulu-2 sten-1 gert-1 kim-1 lone-1 57°00' 4°00' 6°00' 56°00' 25 km 250 km c offee soil fault north sea denmark n s uk g nl p sir i c an yo n mid north sea high central graben a b norwegian–danish basin c entral g raben 100 km north polish strait east shetland platform fennoscandian shield scottish high v ik in g g ra be n rhenish massif bohemian massif jylland sjælland storebælt moray firth mid north sea high ringkøbing–fynhigh fig. 1. location maps showing the position of wells used in the study (a) and major structural elements in the greater north sea area (b) mentioned in the text. on the well map (a) are indicated the locations of the seismic sections shown in figs 49, 50, 56, 58 and 61. grey shading on this map indicates the margins of the siri canyon; grey shading inside the canyon indicates an area of positive relief within the canyon. ggggg, germany; nnnnn, norway; nlnlnlnlnl, netherlands; ppppp, poland; sssss, sweden; u ku ku ku ku k, united kingdom. 7 introduction intense drilling activity following the discovery of the siri field in 1995 has resulted in an improved understanding ofthesiliciclasticpalaeogenesedimentpackageinthedanish sector of the north sea (fig. 1). many of the new wells were drilled in the search for oil reservoirs in sandstone bodies of paleocene–eocene age. the existing lithostratigraphy was established on the basis of data from a generation of wells that were drilled with deeper stratigraphic targets, with little or no interest in the overlying palaeogene sedimentary succession. this means that this early scheme does not include palaeogene sandstone units in the danish sector. in order to improve the understanding of the distribution, morphology and age of the palaeogene sediments, in particular the economically important sandstone bodies, a detailed study of this succession in the danish sector has been carried out. the main aim was to update the lithostratigraphic framework of the succession on the basis of new data from recently drilled wells. all of the widespread palaeogene mudstone units in the north sea were established with norwegian or united kingdom (uk) type wells. in the present work, these units have been maintained unchanged or with only slight modifications. danish reference wells have been established for the units, however, and lithological descriptions have been expanded to cover the characteristics of these units in the danish sector. many of the sandstone bodies recently discovered in the danish sector have a limited spatial distribution and are derived from sources different from those of most of the contemporaneous sandstone bodies in the norwegian and uk sectors; furthermore, the danish sandstone bodies probably neither overlap nor are in contact with the norwegian/uk sandstones. these units have therefore been established as new in the danish sector, and have been assigned danish type and reference sections. the lithostratigraphy presented herein (fig. 2) has its base at the top of the early paleocene (danian) ekofisk formation (chalk group). the top of the study section is at the unconformity between the late eocene – mid-miocene westray group and the mid-miocene to recent nordland group. oil companies operating in the north sea have collected a substantial amount of lithostratigraphic data on the palaeogene successions and a detailed lithostratigraphy has been developed for the danish and norwegian sectors (see e.g. hamberg et al. 2005). a number of informal lithostratigraphic units have been introduced that have subsequently found their way into academia and geological survey organisations. it has been the aim of the present work formally to define these new units. this has been done maintaining their original (albeit informal) names whenever feasible. it has not been the aim of this work to provide a sequence stratigraphic model for the palaeogene sediments in the central and eastern north sea; for this the reader is referred to michelsen et al. (1992, 1995, 1998), mudge & bujak (1994, 1996a, b), neal et al. (1994) and danielsen et al. (1997). the present contribution does not attempt to review the petroleum-related aspects of the palaeogene succession. information about this may be found elsewhere, for example in the annual reports from the danish energy authority. preliminary results from the present work, including a revised lithostratigraphic scheme, were previously published in a brief review paper (schiøler et al. 2005). the present contribution formally describes the new stratigraphic units suggested in the review paper and further documents the palaeogene – lower neogene lithostratigraphy in the danish sector of the north sea. 8 danian selandian thanetian sparnacian ypresian lutetian ma m id dl e eo ce ne lo w er e oc en e u pp er p al eo ce ne c ha lk g ro up r og al an d g ro up st ro ns ay g ro up w es tr ay g ro up st ro ns ay g ro up lo w er p al eo ce ne priabonian rupelian chattian aquitanian burdigalian langhian serravallian 15 20 25 30 35 45 50 55 60 65 m id dl e m io ce ne lo w er m io ce ne lo w er n eo ge ne pa la eo ge ne pa la eo ge ne u pp er o lig oc en e lo w er o lig oc en e u pp er e oc en e m id dl e eo ce ne lillebælt clay fm horda fm røsnæs clay fm fur fm f ølst fm stolle klint clay østerrende clay holmehus fm ve mb bue mb rind mb æbelø fm danian limestone ekofisk fm våle fm li st a fo rm at io n vile mb sele fm balder fm bartonian lutetian 40 søvind marl fm søvind marl fm viborg fm linde clay branden clay vejle fjord fm hodde fm gram fm lark fm (undivided) horda fm se ri es sy st em st ag e d en m ar k on sh or e danish north sea kolga mb hefring mb tyr mb bor mb 49.0 55.5 54.5 57.9 60.0 41.3 37.0 33.7 28.5 23.8 20.5 16.4 14.8 dufa mb odderup fm bastrup sand ribe fm freja mb idun mb n or dl an d g ro up arnum fm kerteminde marl lellinge greensand 9 geological setting the danish sector of the north sea is situated in the central and eastern north sea and comprises three major structural elements: the central graben, the norwegian–danish basin (the eastern part of the northern north sea basin of rhys 1974) and the ringkøbing–fyn high (fig. 1; the geographic terminology and names of structural elements in the north sea used herein are adapted from rhys 1974, rønnevik et al. 1975, deegan & scull 1977 and fyfe et al. 2003). the western boundary of the danish sector largely coincides with the eastern boundary of the mid north sea high, the southern boundary largely coincides with the southern limit of the ringkøbing–fyn high, and the northern boundary is in the norwegian– danish basin. this basin as well as the ringkøbing–fyn high are early permian structures. active rifting occurred in the central graben from the middle to late jurassic along pre-established palaeozoic fault trends. major tectonic activity around the palaeozoic and jurassic structures had largely ceased by late cretaceous time, and the sediment basin below the central north sea was largely characterised by regional subsidence (ziegler 1981). during the late cretaceous to danian sea-level high, pelagic chalk sediments draped the structural highs and the northern and southern north sea basins became one north sea basin delimited by the fennoscandian shield to the north-east, the rheinish–bohemian massif to the south and the british massifs, highs and platforms to the west (see ziegler 1981 fig. 16 for details). chalk sedimentation continued through to the end of the danian stage when it gave way to hemipelagic and siliciclastic sedimentation. this was probably caused by uplift of the basin margins to the west and east (ahmadi et al. 2003). however, most of the siliciclastic sediments were derived from the scottish high and the east shetland platform, uplifted by the iceland plume (ahmadi et al. 2003). by the time of peak uplift, in the mid-thanetian, large sand systems were building out towards the central north sea. most sediment came from the west, but the siri canyon system, a depression in the top chalk surface, was fed from the fennoscandian shield in the north-east and north (fig 1; ahmadi et al. 2003; hamberg et al. 2005). thermal subsidence centered above the central graben continued through the eocene as sea level fell and the temperature decreased. shallow-marine sediments characterised the margins of the north sea basin, especially its western margin, whereas basinal mudstone continued to accumulate in the basin centre and in the eastern part of the basin (joy 1996). inversions controlled by compression between the atlantic spreading zone to the northwest and the orogenesis of the alps to the south added to further uplift of the basin margins and submarine fans and turbidites were deposited near the centre of the basin (jones et al. 2003). during the oligocene, the north sea basin became part of a larger nw european basin. connection with the north atlantic broadened and enhanced communication with the oceanic water mass to the north-west, whereas the connection to the south through the north polish strait became closed for the deep water (fyfe et al. 2003). glacio-eustatic sea-level changes became more frequent and controlled the sedimentary cycles. the eastward progradation direction of the paleocene and eocene sediments gave way to sediment supply from the european massifs to the far south (fyfe et al. 2003). continued subsidence above the mesozoic rift structures created accommodation space for thick sediment packages of basinal mudstones, and few sandstone units reached the basin depocentre above the mesozoic rifts (fyfe et al. 2003). in the neogene epoch, sediment started to be derived from the fennoscandian shield to the north, and the progradation direction changed to the south-west and west in the danish sector of the north sea. facing page: fig. 2. lithostratigraphic column for the palaeogene and lower neogene of the danish north sea sector showing the approximate correlation with danish onshore stratigraphic units. timescale from hardenbol et al. (1998), except for the age of the paleocene–eocene boundary, which is adapted from berggren & aubry (1996) and the age of the sparnacian–ypresian boundary, which is from aubry et al. (2003). stratigraphy and ages of pre-chatian onshore lithostratigraphic units are based on heilmann-clausen (1995) and clemmensen & thomsen (2005). post-rupelian onshore stratigraphy and ages are from dybkjær & rasmussen (2000) and rasmussen (2004a). f, fur formation. 10 previous work the permian to recent lithostratigraphy of the north sea was described in two pioneering stratigraphic works. rhys (1974) provided an overview of the structural elements of the north sea and gave a brief description of the palaeogene sediments. deegan & scull (1977) compiled a detailed lithostratigraphic subdivision and lithological description for the central and northern north sea (figs 3, 4). they subdivided the siliciclastic palaeogene, neogene and quaternary sediments into five major groups: the montrose, moray, rogaland, hordaland and nordland groups. the montrose and moray groups established for the outer moray firth – forties area are proximal equivalents to the rogaland group and are not present in the danish sector, whereas the rogaland, hordaland and nordland groups have widespread distribution in the danish sector. the succession of major mudstone formations contained within the three basinwide groups has formed the backbone of all subsequent lithostratigraphic schemes for the central and northern north sea, including that of the present contribution. the post-danian cainozoic succession of the danish central graben was divided into seven informal units by kristoffersen & bang (1982). the palaeogene comprised five units: north sea marl and cen-1–4 (fig. 4). the ranks of the units were not stated. although descriptions and interpretation of the cen units were detailed, they are essentially informal and have been little used. a revised lithostratigraphy for the palaeogene and neogene of the norwegian north sea sector was published by hardt et al. (1989). their lithostratigraphic scheme includes a number of new palaeogene and neogene sandstone bodies observed in the norwegian and british sectors of the north sea (fig. 4). some of the names of the new sandstone units established by hardt et al. (1989) were subsequently used informally for comparable sandstone units discovered in the danish sector. mudge & copestake (1992a, b) presented a revised palaeogene stratigraphy for the outer moray firth and northern north sea basins. in their papers they redefined the moray and montrose groups of deegan & scull (1977) and abandoned the rogaland group. the authors also demoted the previously established sandstone formations within the two former groups to the rank of members. besides, in an innovative approach they allowed for a greater influence of biostratigraphic data on the characterisation of the various lithostratigraphic units, an approach which is also followed herein. knox & holloway (1992) updated the lithostratigraphic scheme for the palaeogene in the british and norwegian central and northern north sea (figs 3, 4). the authors followed mudge & copestake (1992a, b) in abandoning the rogaland group of deegan & scull (1977), and used mudge & copestake’s revised definition of the montrose and moray groups for the central north sea as well.furthermore, thethickand hitherto undivided hordaland group was subdivided into two new groups, the stronsay group succeeded by the westray group, each containing a distal and a proximal formation. the two distal formations of the two groups, the horda and lark formations, together constitute the bulk of the palaeogene sediments in the danish sector of the north sea and are adopted herein (figs 2–4). although sandstone units occur in both the horda and lark formations in the danish sector, the two proximal sandstone formations of the fig. 3. correlation chart showing the approximate correlation between key lithostratigraphic schemes for the central and eastern north sea at group and formation levels. fu r deegan & scull (1977) hardt et al. (1989) knox & holloway (1992) this study chalk group chalk group chalk group hordaland group lista unnamed unit/ våle sele balder balder balder lista maureen montrose group moray group lark m ou sa sk ad e horda sele rogaland group rogaland group stronsay group westray group stronsay group westray group lista sele horda lark våle nordland group nordland group nordland group 11 stronsay and westray groups, the mousa and skade formations, are absent from the danish sector. following detailed analysis of new, high-resolution seismic surveys covering the succession in the eastern north sea area, efforts were focused on establishing a sequence stratigraphic subdivision of the palaeogene–neogene sediment package. the sedimentary succession was interpreted in a series of publications from a working group at the university of aarhus (e.g. michelsen et al. 1992, 1995, 1998; michelsen 1993; danielsen et al. 1997; huuse & clausen 2001). the result of that work was a subdivision of the palaeogene to mid-neogene sediment package covered by the present work into six genetic units (fig. 4). the sequence stratigraphy of the upper oligocene to miocene in the eastern north sea was dealt with by rasmussen (2004b). further sequence stratigraphic contributions covering the larger north sea basin including the british and norwegian sectors are given by armentrout et al. (1993), mudge & bujak (1994, 1996a, b) and neal et al. (1994). coastal onlap basinward 7 michelsen et al. (1998) 6 6.3 6.2 6.1 5.4 5.3 5.2 5.1 4.4 4.3 4.2 4.1 1.2 1.1 5 4 3 2 1 nordland group hordaland group balder sele lista unnamed unit ekofisk cen-5 cen-4 cen-3 cen-2 cen-1 north sea marl chalk-6 nordland group hordaland group balder sele sele lista lista våle ekofisk nordland group lark horda balder tay m o u sa s ka d e fr ej a k ol ga f u r r in d id un ty r bo r sele lista l is ta fo rt ie s c ro m ar ty m ey v ad e fo rt ie s a n d re w h ei m d al h er m o d f is ke b an k f is ke b an k f ri gg r og al an d g ro u p maureen m au re en ekofisk ekofisk ekofisk nordland group lark horda balder sele bue ve vile vålevåle deegan & scull (1977) northern north sea central north sea kristoffersen & bang (1982) hardt et al. (1989) knox & holloway (1992) this study nordland group balder ty grid ekofisk frigg h ef ri ng d uf a skade fig. 4. correlation chart showing approximate correlation between key lithostratigraphic schemes for the central and eastern north sea and the norwegian part of the northern north sea at formation and member levels. the sequence stratigraphic subdivision of michelsen et al. (1998) is added for comparison. sandstone-dominated units indicated in yellow. 12 material and methods the present lithostratigraphic subdivision represents the combined results from studies of petrophysical logs, biostratigraphyandseismicprofiles,cuttings samples and cored sections. petrophysical logs from c. 70 wells in the danish sector have been scrutinised (see fig. 1 for well locations). the wells have been correlated using petrophysical logs, predominantly gamma-ray and sonic logs. five log panels form the basis for the log correlation (plates 1–5). lithostratigraphic well correlation has been supported by biostratigraphic data: biostratigraphic reports from 29 wells have been re-assessed with the aim of identifying key micropalaeontological and palynological events that occur consistently within the study area (taxa used are planktonic and benthic foraminifers, diatoms, radiolaria, sporomorphs and dinoflagellate cysts). moreover, biostratigraphic sample suites from 11 north sea wells have been prepared at the geological survey of denmark and greenland in order to further determine the biostratigraphic event succession. the bulk of material studied for biostratigraphy is based on cuttings samples, and only few table 1. well data for the new type and reference wells in the danish sector of the north sea augusta-1 cecilie-1 cleo-1 connie-1 e-8 f-1 floki-1 francisca-1 frida-1 inez-1 k-1 mona-1 nini-3 sandra-1 siri-1 siri-2 siri-3 tabita-1 bor mb(t), bue mb(t),ve mb(t) bor mb(r), tyr mb(r) bue mb(r), lista fm(r), ve mb(r), vile mb(r) idun mb(t), rind mb(t) bue mb(r), lista fm(r), ve mb(r), vile mb(r), våle fm(r) dufa mb(r) hefring mb(t) freja mb(t) freja mb(r) dufa mb(t), fur fm(r) fur fm(r) balder fm(r), horda fm(r), lark fm(r) kolga mb(r), tyr mb(t) rind mb(r) horda fm(r), lark fm(r), sele fm(r), våle fm(r) idun mb(r) balder fm(r), kolga mb(t), vile mb(t) sele fm(r) 56°17´57.40´́ n 04°24´04.64´́ e 56°24´23.73´́ n 04°45´42.00´́ e 56°23´23.54´́ n 04°25´22.70´́ e 56°24´28.34´́ n 04°42´30.36´́ e 55°38´13.42´́ n 04°59´11.96´́ e 57°01´53.4´́ n 06°54´28.6´́ e 56°27´48.58´́ n 05°16´47.11´́ e 56°22´27.95´́ n 04°48´05.30´́ e 56°17´14.15´́ n 05°01´50.20´́ e 56°50´28.39´́ n 06°57´41.62´́ e 57°07´37.74´́ n 07°09´43.11´́ e 56°16´35.94´́ n 04°00´15.81´́ e 56°41´31.96´́ n 05°24´12.35´́ e 56°35´13.33´́ n 05°01´35.19´́ e 56°29´11.10´́ n 04°54´57.49´́ e 56°29´40.53´́ n 04°52´13.26´́ e 56°30´34.92´́ n 05°03´48.27´́ e 56°13´37.50´́ n 04°23´47.56´́ e 04.03.2001 2991.0 mdrt 37.8 rt 65 15.10.2000 2361.0 mdrt 37.8 rt 59.4 06.02.1984 4866.1 mdkb 40.5 kb 63.1 02.02.2001 2351.8 mdrt 37.8 rt 61.5 08.04.1994 2527.4 mdkb 36.6 kb 43.6 06.10.1968 2421.6 mdkb 37.19 kb 40.8 29.08.2000 1878 mdrt 35.8 rt 53.2 20.07.1998 1888.5 mdrt 36.4 kb 60 26.07.1997 2274 mdrt 39.0 rt 54.3 11.09.1977 1983.9 mdkb 35.1 kb 35.4 22.01.1970 2292.4 mdkb 37.2 kb 56.4 03.10.1982 4241.6 mdkb 36.6 kb 65.5 12.01.2001 1851.2 mdrt 37.3 rt 58.2 18.06.1998 2139 mdrt 36 kb 65 28.11.1995 2220 mdkb 23 kb 60 03.08.1996 2297.5 mdrt 36.6 rt 60.6 30.08.1996 2171.5 mdrt 36.6 rt 60.1 10.09.1983 4353 mdkb 40 kb 65 dong e&p a/s dong e&p a/s chevron petroleum co. dong e&p a/s maersk oil & gas a/s gulf oil company kerr-mcgee int. aps dansk operatørselskab i/s dansk operatørselskab i/s chevron petroleum co. california oil co. chevron petroleum co. dong e&p a/s statoil e&p a/s statoil e&p a/s statoil e&p a/s statoil e&p a/s statoil e&p a/s type (t) or reference (r) well coordinates operator spud date td (logger’s kb/rt elevation water for listed units: depth in m) (m above msl) depth (m)well fm: formation. mb: member. mdrt: measured depth below rotary table. mdkb: measured depth below kelly bushing. 13 fig. 5. chronostratigraphy and biostratigraphy of the paleocene – middle miocene. a: paleocene–eocene. b: eocene–oligocene. c: oligocene – middle miocene. calibration of chronostratigraphic units follows hardenbol et al. (1998), berggren & aubry (1996) for the paleocene–eocene boundary and aubry et al. (2003) for the sparnacian–ypresian boundary. key dinoflagellate datums are calibrated mainly using age estimates from hardenbol et al. (1998) and williams et al. (2004). key microfossil datums are calibrated via their correlation with calibrated dinoflagellate datums as suggested by mudge & bujak (1996b), using age estimates from hardenbol et al. (1998) and williams et al. (2004). the combined event succession is correlated with the north sea microfossil zonation of king (1989) and lithostratigraphic units treated herein. in the microfossil event column, the planktonic foraminifer events appear in normal font, benthic foraminifers in italics; diatoms and radiolarians are underlined. senoniasphaera inornata palynodinium grallator, dinogymnium spp. alisocysta reticulata abundant p. pyrophorum isabelidinium? viborgense p. pyrophorum, p. australinum acme a. gippingensis alisocysta margarita common cerodinium wardenense apectodinium augustum apectodinium augustum, acme apectodinium spp. acme d. oebisfeldensis, influx inaperturopollenites spp., common h. tubiferum deflandrea oebisfeldensis dracodinium condylos nsp6 (pars) nsp5b nsp5a nsp4 nsp3 nsp2 nsp1 a b c nsb4 (pars) nsb3b nsb3a nsb2 nsb1 a b c horda balder sele ekofisk tor våle bue ve vile lista planktonic foraminifers benthic foraminifers diatoms and radiolaria planktonic microfossils north sea biozones (king 1989) lithostratigraphy selected biostratigraphic events used in the present studygeochronology ma a chronostratigraphy (berggren et al. 1995) c re ta ce ou s (p ar s) pa le oc en e eo ce ne (p ar s) u pp er (p ar s) lo w er u pp er lo w er (p ar s) fm mbdinoflagellate cysts ypresian (pars) maastrictian (pars) thanetian sparnacian selandian danian 54.5 55.5 57.9 60.0 65.0 50 55 60 65 uvigerina batjesi turrillina brevispira gaudryina hiltermanni common subbotina ex gr. linaperta fenestrella antiqua, foraminifers very rare impoverished benthic agglutinated assemblage common globoconusa daubjergensis globanomalina cf. compressa, s. trivialis increasing diversity of calcareous foraminifers reappearance of planktonic foraminifers increasing diversity of calcareous benthic foraminifers cenodiscus spp., cenosphaera spp. cretaceous foraminifers common f. antiqua and coscinodiscus morsianus pseudotextularia elegans benthic microfossilsstageseries 14 35 40 50 45 lark formation rupelian (pars) o lig oc en e (p ar s) eo ce ne ( pa rs ) priabonian 41.3 bartonian lutetian ypresian (pars) nsb7a nsb6b nsb6a nsb5c nsb5b nsb5a nsb4 nsb3a nsb2 (pars) nsb3b planulina costata pseudohastigerina spp. abundant radiolaria (cenosphaera spp.), cyclammina amplectens lenticulina gutticostata, spiroplectammina spectabilis balder horda eatonicysta ursulae diphyes ficusoides areosphaeridium michoudii heteraulacacysta porosa diphyes colligerum areosphaeridium diktyoplokum common e. ursulae phthanoperidinium clithridium globigerinatheka index cibicidoides truncanus vaginulinopsis decorata 49.0 37.0 33.7 uvigerina batjesi turrillina brevispira gaudryina hiltermanni common subbotina patagonica dracodinium condylos deflandrea oebisfeldensis acme d. oebisfeldensis, influx inaperturopollenites spp., common h. tubiferum fenestrella antiqua, foraminifers very rare cerebrocysta bartonensis uvigerina germanica karrulina conversa corrudinium incompositum sele (pars) nsp9b nsp9a nsp8c nsp8b nsp8a nsp7 nsp6 nsp5b nsp4 (pars) nsp5a lo w er ( pa rs ) u pp er m id dl e lo w er ( pa rs ) planktonic foraminifers benthic foraminifers diatoms and radiolaria planktonic microfossils north sea biozones (king 1989) lithostratigraphy selected biostratigraphic events used in the present study chronostratigraphy (berggren et al. 1995) fm mbdinoflagellate cysts benthic microfossilsstageseries geochronology ma b fig. 5b. chronostratigraphy and biostratigraphy of the eocene–oligocene. 15 lark nsp9a (pars) nsb6b (pars) uvigerina germanica karrulina conversa nsp14b nsb13a nsp14a nsb12c nsp13 nsb12b nsb12a nsp12 nsb11 nsp11 nsp10 nsb10 nsb9 nsp9c nsb8c nsb8b nsb8a nsb7b nsb7a nsp9b aulacodiscus allorgei turrillina alsatica bolboforma spiralis asterigerina staeschei, elphidium inflatum, meonis pompilioides uvigerina tenuipustulata plectofrondicularia seminuda aulacodiscus insignis quadrata (small), b. antiqua, g. girardana common elphidium subnodosum, common paragloborotalia nana rotaliatina bulimoides “turborotalia” ampliapertura common a. guerichi, paragloborotalia opima s.s. bolboforma metzmacheri pararotalia canui aulacodiscus insignis quadrata (large) spirosigmoilinella compressa cibicidoides mexicanus gyroidina mamillata wetzeliella gochtii phthanoperidinium amoenum chiropteridium spp. membranophoridium aspinatum distatodinium biffi cordosphaeridium cantharellus apteodinium spiridoides caligodinium amiculum thalassiphora pelagica hystrichokolpoma cinctum rhombodinium draco corrudinium incompositum achilleodinium biformoides enneadocysta pectiniformis burdigalian aquitanian m io ce ne (p ar s) lo w er chattian u pp er rupelian (pars) lo w er (p ar s)o lig oc en e (p ar s) m id dl e langhian serravallian 28.5 23.8 20.5 16.4 14.8 11.2 cousteaudinium aubryae nordland group bulimina elongata bolboforma clodiusi p. comatum 15 20 30 25 cannosphaeropsis passio tortonian (pars)u pp er (p ar s) planktonic foraminifers benthic foraminifers diatoms and radiolaria planktonic microfossils north sea biozones (king 1989) lithostratigraphy selected biostratigraphic events used in the present study chronostratigraphy (berggren et al. 1995) fm mbdinoflagellate cysts benthic microfossilsstageseries geochronology ma c fig 5c. chronostratigraphy and biostratigraphy of the oligocene – middle miocene. 16 core samples have been available. as the use of stratigraphic lowest occurrences (lo) of taxa in cuttings samples may be hampered due to downhole caving, the event succession comprises almost exclusively stratigraphic highest occurrences (ho) of taxa (a single significant lo is included in the succession). the event succession is shown in fig. 5a–c; its correlation with international and north sea biozones is shown in fig. 6a–c. seismic sections from the 2-d and 3-d seismic surveys cgd85, dk-1, rtd81–re94, ucg96 and ucge97 have been used to further support the well correlation and to map the stratigraphic units in areas with only scattered well coverage. the combined results from the correlation and mapping procedures are presented as isochore maps for individual stratigraphic units. inspection of cuttings samples from 16 key wells supplemented with sedimentological studies of cored intervals from 23 wells have formed the basis for the lithological and sedimentological descriptions of the units. the well depths mentioned in the lithostratigraphy section are loggers’ depths measured either from rotary table (mdrt) or kelly bushing (mdkb). supplementary data for new type and reference wells are provided in table 1. the names assigned to the new lithostratigraphic units definedhereinare derived from nordic mythology and thus follow the nomenclatural tradition previously established for the norwegian north sea (isaksen & tonstad 1989). it should be noted that the micropalaeontology-based palaeoenvironmental terminology used herein was originally developed for a passive margin situation (e.g. the terms ‘neritic’ and ‘bathyal’ to indicate the physiographic zones ‘shelf ’ and ‘shelfslope’, respectively). its application herein to the epicontinental north sea basin solely relates to depositional depth. offshore and onshore lithostratigraphic nomenclature there is a high degree of lithological similarity between the palaeogene–neogene mudstone succession in danish offshore boreholes and that in onshore exposures and boreholes. however, the status of the danish onshore units is quite varied since many units were named before a standard for description of a lithostratigraphic unit was established; some fulfil these requirements, whereas others are still informal. if a previously established onshore unit and an offshore unit can be demonstrated to be identical (e.g. the holmehus formation and the new ve member proposed herein), the name of the onshore unit theoretically has priority over the name of the offshore unit (salvador 1994). in other cases, names of offshore units can be argued to have priority over onshore units (e.g. sele and balder formations over ølst formation). however, in order to acknowledge the traditional distinction between offshore and onshore stratigraphic nomenclature, the two sets of nomenclature are kept separate herein. whenever possible, comments are given in the text to explain the relationship between offshore and onshore danish stratigraphic nomenclature. a correlation between the two sets of nomenclature is shown in fig. 2. chronostratigraphy and biostratigraphy age assessment of the lithostratigraphic units in the north sea sedimentary succession is based on correlation between key biostratigraphic events encountered in the units and the calibrated standard chronostratigraphy published by berggren et al. (1995), with modification for the paleocene–eocene boundary following ratification of its position by the international union of geological scientists (aubry et al. 2002). the key events are from biostratigraphic zonation schemes established for the north sea area. planktonic and benthic microfossils are covered by the zonation schemes of king (1983, 1989; figs 5a–c, 6a–c). dinoflagellates from the paleocene and eocene epochs are covered by the zonation scheme of mudge & bujak (1996b; fig. 6a, b); the oligocene and miocene epochs are covered by the zonation schemes of costa & manum (1988) with modifications by köthe (1990, 2003; fig. 6b, c). key events from these schemes used in this study are listed in fig. 5a–c. for the dinoflagellate events, geochronological calibration has been largely established using age estimates from hardenbol et al. (1998), munsterman & brinkhuis (2004) and williams et al. (2004). for events not mentioned in these three publications, the works of mudge & bujak 17 p2 p9 p7 p6 b a p5 p4 c p8 np13 np12 np10 np9 np11 b a b a p3 c b a p1 pα + p0 np8 np6 np5 np4 np3 np2 np1 np7 abathomphalus mayaroensis cc26 cc25 (pars) pseudotextularia elegans p6 p5 p4 p3 p2 p1 e1a e2a e2b e2c e3a e3b l e1c e1beo ce ne (p ar s) pa le oc en e c re ta ce ou s (p ar s) np14 (pars) e3c 50 60 55 65 nsp6 (pars) nsp5b nsp4 nsb4 (pars) nsb3a nsb2 nsp5a nsb3b nsp3 nsp2 nsp1 nsb1 65.0 ypresian (pars) lo w er (p ar s) u pp er lo w er 55.5 60.0 thanetian selandian danian u pp er (p ar s) maastrichtian (pars) 57.9 54.5 sparnacian c b a b c a planktonic microfossils benthic microfossils dinoflagellate cystsplanktonic microfossils calcareous nannofossils north sea biozonesstandard biozones chronostratigraphy (berggren et al. 1995) stageseries berggren & miller (1988), berggren et al. (1995) mudge & bujak (1996b) martini (1971) king (1989) geochronology ma a fig. 6. biostratigraphic correlation charts showing approximate correlation of calibrated standard planktonic foraminifer and nannofossil biozones with north sea microfossil and dinoflagellate biozones. calibration of the standard biozones follows hardenbol et al. (1998). relationships between the north sea biozones are approximate and their correlation with the standard zones may deviate from that of the original authors (for discussion, see text). a: paleocene–eocene biostratigraphic correlation chart. b: eocene–oligocene biostratigraphic correlation chart. c: oligocene – middle miocene biostratigraphic correlation chart. p17 e6b np14 p9 p7 p8 np13 np12 np11 e2a e2b e2c e3a e3b e1c e1b e3c nsp6 nsp5b nsb4 nsb3ansp5a nsb3b p18 p16 p15 np23 (pars) np22 np21 np19–20 np18 p14 p12 p11 p10 p13 np17 np16 np15 e3d e4a e4b e4c e4d e5a e5b e6a e6c e7a e7b e8b e8a d13 mudge & bujak (1996b) costa & manum (1988), köthe (1990) planktonic microfossils benthic microfossils dinoflagellate cystsplanktonic microfossils calcareous nannofossils north sea biozonesstandard biozones p19 (pars) 35 40 50 45 o lig oc en e (p ar s) lo w er ( pa rs ) eo ce ne (p ar s) u pp er m id dl e lo w er ( pa rs ) nsb7ansp9b nsp9a nsb6b nsp8c nsb6a nsp8b nsb5c nsp8a nsp7 nsb5b nsb5a p6 b a np10 (pars) nsp4 (pars) nsb2 (pars) e1a (pars) geochronology ma chronostratigraphy (berggren et al. 1995) stageseries berggren & miller (1988), berggren et al. (1995) costa & manum (1988), köthe (1990), mudge & bujak (1996b) martini (1971) king (1989) rupelian (pars) priabonian 41.3 lutetian ypresian (pars) 49.0 37.0 33.7 bartonian b fig. 6b. eocene–oligocene biostratigraphic correlation chart. 18 m7 m9 m12 m8 m11m10 m6 m5 m3 m2 m4 m1 b a p22 p21 b a p20 p19 p18 nn5 nn6 nn9a– nn7 nn4 nn3 nn2 nn1 np25 np24 np23 np22 np21 (pars) d13 d14 d15 d16 d17 d18 d19 tortonian (pars) burdigalian aquitanian m io ce ne (p ar s) lo w er chattian u pp er rupelian (pars) lo w er ( pa rs ) o lig oc en e (p ar s) m id dl e langhian serravallian 28.5 23.8 20.5 16.4 14.8 11.2 15 20 30 25 u pp er (p ar s) nsp9a (pars) nsb6b (pars) nsp14b nsb13a nsp14a nsb12c nsp13 nsb12b nsb12a nsp12 nsb11 nsp11 nsp10 nsb10 nsb9 nsp9c nsb8c nsb8b nsb8a nsb7b nsb7a nsp9b m13a (pars) nn9b (pars) planktonic microfossils benthic microfossils dinoflagellate cystsplanktonic microfossils calcareous nannofossils north sea biozonesstandard biozones chronostratigraphy (berggren et al. 1995) stageseries berggren & miller (1988), berggren et al. (1995) costa & manum (1988), köthe (1990), martini (1971) king (1989) geochronology ma c fig. 6c. oligocene – middle miocene biostratigraphic correlation chart. 19 20 (1996b), dybkjær (2004), piasecki (2005) and schiøler (2005) have been consulted. however, whereas hardenbol et al. (1998) and williams et al. (2004) used the timescale of berggren et al. (1995), mudge & bujak used the slightly older timescale from haq et al. (1987) for calibration of their events. therefore, the ages of events only listed by mudge & bujak have been recalibrated herein to conform to the timescale of berggren et al. (1995). king (1989) calibrated his planktonic and benthic microfossil zone markers with the standard chronostratigraphic scale of berggren et al. (1985a, b). however, king noted that only a few first-order correlations were possible; most of the calibrations were made using dinoflagellates, planktonic foraminifers and nannoplankton from onshore sections in the north sea basin (king 1989 p. 420); the correlation of the lower miocene is particularly uncertain (king 1989 p. 446). paleocene and eocene key planktonic and benthic microfossil events from king (1989) were subsequently correlated with the north sea dinoflagellate events by mudge & bujak (1996b). by using the above-mentioned recalibration of key dinoflagellate events from mudge & bujak (1996b), it is feasible to indirectly correlate king’s north sea microfossil events with the timescale of berggren et al. (1995). this has been attempted in fig. 5a–c. figure 6a–c shows the relationships between the north sea biozones and their correlation with the standard planktonic foraminifer and calcareous nannofossil zones. however, it should be noticed that in a few cases the correlation of the north sea microfossil and dinoflagellate zones with the standard zones in fig. 6a–c is at variance with that of the authors of the same zones. this is an effect of improved age determinations of the standard zones and the dinoflagellate events used to calibrate the north sea microfossil zones. the section below outlines the current status for the palaeogeneandneogenechronostratigraphicunitscovered bythestudiedsuccessionandlistskeybiostratigraphic events used for chronostratigraphic correlation of the succession. paleocene the bases of the selandian and thanetian stages, which together constitute the upper paleocene series, have yet to be formally defined. however, ongoing work in the international subcommission on palaeogene stratigraphy indicates that the global standard stratotype-section and point (gssp) of the base of the selandian stage will probably be close to the p2–p3a or the p3a–p3b standard planktonic foraminifer zone boundary, while the gssp for the thanetian stage will probably be at the base of magnetochron c26n (gradstein & ogg 2002). hardenbol et al. (1998) followed berggren et al. (1995) in placing the base of the selandian stage at the base of zone p3a, at the lowest occurrence of the planktonic foraminifer morozovella angulata. however, many of the microfossil species that characterise the danian–selandian boundary interval in the international zonation schemes, including m. angulata, are extremely rare or absent in the north sea basin thereby hampering chronostratigraphic correlation of the boundary. based on a study of core material from the type area for the danian and selandian stages, clemmensen & thomsen (2005) concluded that the danian–selandian stage boundary is located in the upper part of the np4 standard nannofossil zone, close to the np4–np5 zone boundary, approximately at the p3a– p3b zone boundary, at c. 60 ma on the timescale of hardenbol et al. (1998). they further concluded that there is a hiatus between the danian and selandian stages in the danish area outside the central graben due to truncation of the danian limestones of the ekofisk formation (fig. 5a; clemmensen & thomsen 2005). hence, the danian–selandian stage boundary is herein placed just below the downhole reappearance (provisional ho) of planktonic foraminifers and the ho of the dinoflagellate alisocysta reticulata, but above the closely spaced events marked by the ho of the planktonic foraminifers subbotina trivialis and globanomalina cf. compressa (e.g. jones 1999; mudge & bujak 2001). the selandian–thanetian stage boundary is herein approximated by the ho of the dinoflagellate palaeoperidinium pyrophorum, at the base of the p5 dinoflagellate zone of mudge & bujak (1996b). this level is close to the base of magnetochron c26n, according to hardenbol et al. (1998). eocene the base of the eocene is at the base of the negative carbon isotope excursion (cie) at 55.5 ma (berggren & aubry 1996; aubry et al. 2002). this position is below the base of the ypresian stage, the lowermost eocene stage. therefore it has been proposed to reintroduce the sparnacian stage as the new basal eocene stage between the cie and the base of the ypresian (aubry et al. 2003). the cie has been correlated with the proliferation of the dinoflagellate genus apectodinium, an event recognised globally (e.g. knox 1996; crouch et al. 2001). onshore denmark, the cie and the proliferation of apectodinium coincides precisely with the laminated stolle klint clay in the 21 2900 m 3000 2900 m 3000 2000 m 2100 2700 m 2900 m horda fm balder fm sele fm lista fm bue mb r og al an d g ro up st ro ns ay g ro up ve mb vile mb våle fm chalk group 3000 2800 kim-1 gr sonic gr sonic gr sonic gr sonic gr sonic kim-1 mona-1 cleo-1 gulnare-1 e-8 e-8 lowermost part of the haslund member of the ølst formation (heilmann-clausen & schmitz 2000; willumsen 2004). in the north sea basin, the acme of apectodinium is located in the lowermost, laminated part of the sele formation (sensu deegan & scull 1977, see below) according to knox (1996). as the event is a lo, its position cannot be determined with certainty in wells in which this interval is covered only by cuttings samples. in the north sea basin, however, this stratigraphic level is characterised by a prominent excursion on the gamma-ray log near the base of the sele formation which therefore can be used as an approximation for the base of the eocene series. the remaining stages of the eocene series, the ypresian, lutetian, bartonian and priabonian stages, lack basal boundary gssps for the present. in this paper, we follow mudge & bujak (1996b) and approximate the bases of the three latter stages by using three key dinoflagellate events: the base of the lutetian stage is at the ho of common eatonicysta ursulae, the base of the bartonian stage is close to the ho of diphyes colligerum, and the base of the priabonian stage is close to the ho of heteraulacacysta porosa. the base of the classic ypresian stage is at the lo of the calcareous nannoplankton species tribrachiatus digitalis. as yet, there is no commonly recognised ho index event at that level in the north sea basin, but the boundary between the sparnacian and the ypresian stages may be placed below the hos of common cerodinium wardenense and apectodinium augustum (fig. 5a), both dinoflagellate species. oligocene the gssp for the eocene–oligocene boundary is in the massignano section (central italy), at the highest occurrence of the planktonic foraminifer genera hantkenina and cribrohantkenina, immediately above the p17–p18 planktonic foraminifer zone boundary (premoli silva & jenkins 1993). however, hantkeninids have not been observed from the north sea basin and alternative zone markers have therefore been used here. in the north sea basin, the planktonic foraminifer globigerinatheka index and the benthic foraminifer cibicidoides truncanus have their hos in the uppermost eocene (king 1989), and the two events may be used to approximate the eocene–oligocene boundary. a palynological marker of the lowermost oligocene is the ho of the dinoflagellate areosphaeridium diktyoplokum (brinkhuis & biffi 1993; brinkhuis & visscher 1995), which is widespread in the north sea basin. the three latter events in combination serve as useful markers for bracketing the eocene–oligocene boundary in the north sea basin. the principal criterion for the rupelian–chattian (lower–upper oligocene) boundary has not yet been defig. 7. log panel illustrating the thickness variation of the rogaland group formations in the danish central graben. 22 cided by the subcommission on palaeogene stratigraphy. indications are that the boundary may be positioned at the base of the p21b planktonic foraminifer zone (premoli silva 2005), at 28.5 ma (hardenbol et al. 1998). however, the defining boundary event cannot be recognised in the north sea basin and its exact correlation with the north sea biostratigraphic event succession remains uncertain. instead, most north sea biostratigraphers recognise the rupelian–chattian stage boundary at the ho of the benthic foraminifer rotaliatina bulimoides. this event marks the top of the nsb7 zone of king (1983, 1989; fig. 5c) and the nsr7 zone of gradstein et al. (1994). the ho of r. bulimoides is at 29 ma in the northern north sea according to gradstein & bäckström (1996), slightly older than the 28.5 ma for the rupelian–chattian stage boundary quoted by hardenbol et al. (1998). the rupelian–chattian stage boundary may also be approximated by the ho of the dinoflagellate rhombodinium draco. in the north sea wells reported herein, where both the hos of r. bulimoides and r. draco have been recorded, these events are largely contemporaneous. however, in the type area of the rupelian and chattian stages, r. draco has its ho above r. bulimoides in the type chattian (van simaeys et al. 2004). therefore, it may be inferred that the two latter events probably bracket the rupelian–chattian boundary (fig. 5c). ve mb tyr mb bue mb log depth core depth våle fm lista fm bor mb gr soniccecilie-1 clay si. vf. f. m. sand c. vc. p p p p p p p p p ? s 5o 2240 2250 2260 2270 2280 2240 2250 2260 2270 2280 vile mb fig. 8. core log showing intrusive sandstones in the våle and lista formations in the cecilie-1 well. for legend, see fig. 9. the two intervals marked by grey bars in the core depth column are shown as core photographs in fig. 10. 23 lithology sedimentary structures sandstone intrusions mudstone clasts chert siderite calcite concretions trace fossils sandstone mudstone marl chalk carbonate cement (non-calcitic) calcite cement pyrite glaucony parallel lamination faint parallel lamination water-escape pipes (large) load cast dish structures and pipes deformed/slumped bedding fractures/faults bed boundary cross-lamination sandstone intrusions flow structures stylolites zoophycos helminthopsis planolites thalassinoides chondrites low moderate intense degree of bioturbation p g s c miocene the oligocene–miocene series boundary is bracketed by a number of hos at its type section (lemme-carosio, north-west italy). unfortunately, none of the foraminifer events are believed to be true stratigraphic tops (facies dependent), and reworking in the section hampers the use of nannofossil tops (steininger et al. 1997). however, the dinoflagellate succession from the lemme-carosio section has been documented in detail by powell (1986), brinkhuis et al. (1992) and zevenboom (1995, 1996), and provides a means for direct correlation to the north sea basin (munsterman & brinkhuis 2004). the ho of distatodinium biffii is below the chattian–aquitanian boundary in its type section and the ho of chiropteridium spp. is above. this succession of events can be recognised in many north sea wells, and the chattian–aquitanian boundary is positioned between the two. supporting microfossil events that characterise the lowermost miocene include the ho of the diatom aulacodiscus insignis quadrata (small morphotype, same as diatom sp. 3 of king 1983, 1989), a widespread event in the north sea basin, and the ho of the benthic foraminifer brizalina antiqua (king 1989). the ho of the planktonic foraminifer paragloborotalia nana marks uppermost chattian strata. the principal criteria for the aquitanian–burdigalian, burdigalian–langhian and langhian–serravallian stage boundaries are as yet undecided. most authors place the three boundaries at microfossil zone boundaries or magnetochron boundaries at 20.5, 16.4 and 14.8 ma, respectively (hardenbol et al. 1998; williams et al. 2004). the correlation of the three boundaries to the north sea basin is feasible using the dinoflagellate zonation scheme of de verteuil & norris (1996), established for us east coast sections and the review of dinoflagellate index events published by williams et al. (2004). the former zonation scheme is correlated directly with the zonation schemes of berggren et al. (1995) and the miocene timescale by means of calcareous nannofossils and foraminifers. the aquitanian–burdigalian boundary is positioned just above the ho of the dinoflagellate caligodinium amiculum. the burdigalian–langhian boundary is placed between the ho of the dinoflagellates hystrichokolpoma cinctum and pyxidinopsis fairhavenensis, two events that bracket the boundary level. the langhian–serravallian boundary is slightly above the ho of the dinoflagellate cousteaudinium aubryae. in this study, these four events have been used to approximate the three stage boundaries. fig. 9. legend for core logs (figs. 8, 11, 18, 27, 30 and 39); the lithological colour scheme is also adopted on well sections (e.g. fig. 13). 24 0 10 20 30 40 50 60 70 80 90 100 cm 2239 m 2240 m 2245 m 2246 m cecilie-1 fig. 10. core photographs showing dark grey, largely structureless, discordant, intrusive sandstones within the lighter grey mudstones of the upper tyr member in the cecilie-1 well. depths are core depths. stratigraphic positions of the figured intervals are shown on fig. 8. lithostratigraphy rogaland group the rogaland group was established by deegan & scull (1977) and comprises the paleocene to lower eocene marlstone and mudstone succession between the top of the ekofisk formation of the chalk group (deegan & scull 1977) and the glaucony-rich mudstones of the hordaland group (now stronsay and westray groups) in the central north sea. in most of the danish sector, the rogaland group has a relatively uniform thickness and comprises the våle, lista, sele and balder formations (fig. 7). the fur formation is a part of the rogaland group and is present in a limited area in the north-eastern part of the danish sector of the north sea stretching into the norwegian sector. hardt et al. (1989) added three new sandstone units to the rogaland group in the southern viking graben (norwegian sector): ty formation, heimdal formation and hermod formation (fig. 4). although these sandstone units are broadly comparable to coeval sandstone units encountered in the siri canyon system (figs 1, 4) in the danish sector, the norwegian units and the siri canyon sandstones have different provenances and are not contiguous with each other. therefore, the sandstone units in the danish sector are described herein as new members. in the siri canyon (fig. 1), the mudstones of the rogaland group contain concordant or discordant postdepositional sandstone intrusions (hamberg et al. 2005). in some wells (e.g. cecilie-1 and nini-3) sandstone intrusions are very common (figs 8–11). the sands have intruded most levels in the rogaland group, but the ve member (new member of the lista formation, see below) in the middle part of the group is particularly rich in intrusions. most of the intrusive sandstones are only a few millimetres thick, but they may reach a thickness of 5 m. in some wells they constitute up to 30% of the total sandstone thickness. the intrusions are usually massive, but faint lamination is locally present, especially at the top of the beds. most of the intrusive sandstone bodies are separated by in situ mudstones, but they may also occur in intervals showing multiple intrusions. the boundaries with the host rock are slightly to very irregular or wavy, and in places discordant. minor intrusive offshoots (apophyses) into the host rock are common. the petrography of the intrusive sandstones is similar to that of the in situ sandstones and they are therefore most likely sourced from the 25 latter. the intrusion of sand was mainly subhorizontal, parallel to the bedding, and most of the intrusions can thus be classified as sills. the sandstone intrusions are either unconsolidated or cemented by calcite or other carbonate minerals. in some wells the intrusive sandstones are chlorite-cemented. in the nini-3 cores, the vile and ve members of the lista formation (new members, see below) are particularly rich in intrusive sandstones (figs 11, 12); the intrusions increase in number and thickness upward through the vile member to terminate in a large intrusion complex in the ve member. commonly, mudstone clasts are abundant in the sandstones and in the intrusion complex in the nini-3 well. where present, they constitute from a few percent up to 90% of the volume of the host sandstones. they are most abundant in the upper parts of the beds. the clasts range in size from a few millimetres to wider than the core diameter. most of the clasts are angular, often with delicate protrusions, and aligned parallel to the bounding planes of the intrusive sandstone body. in thick intrusions, flow banding and dewatering structures are occasionally present. top-bed rip-down mudstone clasts (stow & johansson 2000) are common (fig. 12). fossil wood fragments are present, but rare. as the distribution of intrusive sandstones is the result of postdepositional rather than synsedimentary processes, they may cross lithostratigraphic boundaries. when an injected sandstone body occurs in direct contact with an in situ sandstone unit (e.g. as seen in the higher parts of the new tyr member in fig. 8), it is impossible to distinguish between the two genetically different units on the basis of petrophysical logs and cuttings samples alone; only a sedimentological study of core material may reveal the different nature of the two sandstones. våle formation history. the våle formation was established by hardt et al. (1989) for the marls with interbedded claystones, limestones and siltand sandstone stringers that overlie the chalk group in the central and northern north sea. the lista fm vile mb tyr mb idun mb nini-3 clay si. vf. f. m. sand c. vc. p c c log depth core depthgr sonic 1760 1750 1740 1730 1760 1750 1740 1730 fig. 11. core log showing intrusive sandstones in the vile and idun members in the nini-3 well. for legend, see fig. 9. the two intervals marked by grey bars in the core depth column are shown as core photographs in fig. 12. 26 fig. 12. core photographs of the vile and idun members in the nini3 well showing sandstone intrusions, weak flow banding (1), injection breccia with abundant irregular and angular clasts (2), mudstone clasts in injected sand (3) and top bed rip-down clasts (4). depths are core depths. stratigraphic positions of the figured intervals are shown on fig. 11. fig. 13. e-8, danish reference well for the våle and lista formations, and reference well for the vile, ve and bue members. black bar shows cored section. 1726 m 1727 m nini-3 1739 m 0 10 20 30 40 50 60 70 80 90 100 cm 4 3 1 2 2060.3 2044.0 2030.3 2027.6 2057.0 horda fm balder fm sele fm lista fm bue mb ve mb vile mb våle fm chalk gp 2100 2000 m e-8 gr sonic presence of a marly succession on top of the chalk group was previously noted by deegan & scull (1977) and treated informally as an equivalent to the more coarse-grained maureen formation in the uk sector of the north sea. kristoffersen & bang (1982) established the north sea marl for an exclusively marly and calcareous unit corresponding to the maureen formation-equivalent unit of deegan & scull (1977). although the description of the north sea marl fulfils the requirements for a formal description of a lithostratigraphic unit (with the exception of lacking indication of the rank of the unit), they specifically stated that their unit was only informally established. the name north sea marl has only rarely been used outside the danish sector of the north sea; the sediments are instead referred to the våle formation, which covers most national sectors of the north sea basin. as doubt may be raised about the formal status of the north sea marl unit, and in order to promote communication between north sea stratigraphers, it is considered by the present authors that the våle formation of hardt et al. (1989) serves as the better name for the marlstone unit. type well. norwegian sector well 1/3-1, 3258–3209 m mdkb. danish reference wells. e-8, 2060.3–2057.0 m mdkb (fig. 13);siri-1,2186.5–2156.3mmdkb(fig.14;plates1, 4). distribution and thickness. the våle formation and its equivalents are present throughout the north sea basin, except in a few areas where their absence is due to nondeposition or erosion. the våle formation is absent on 27 pyrite-bearing marlstones dominate the formation (fig. 16). thin sandstone intrusions are present locally. in the siri canyon, the marls are interbedded with turbidite sandstones; where sandstone-dominated, the succession is referred to a new member (bor member, defined below). log characteristics. from its base to its top, the våle formation is characterised by an overall steady increase in gamma-ray response, combined with an overall steady decrease in sonic readings. when the bor member sandstones are present, blocky log signatures with higher gamma-ray values and lower sonic readings interrupt this general trend (fig. 17). boundaries. in most wells in the danish sector, the change from the chalks of the chalk group to the marlstones of the våle formation is gradational and the boundary can be difficult to position (fig. 16). in the siri canyon, however, most wells show an erosional contact between the chalk group and the våle formation and the formation boundary is sharp. on the petrophysical logs, the boundary is placed where the stable, low gamma-ray response characteristic of the ekofisk formation starts to increase upwards and the high sonic readings (also characteristic of the latter formation) start to decrease upwards. the change in the log pattern may be stepwise with each step represented by a small increase in gamma-ray values and an accompanying decrease in sonic readings. the våle formation is overlain by the lista formation. intrabasinal highs (hardt et al. 1989) and in parts of the siri canyon where the rogaland group overlies the chalk group with an erosionalunconformity. its thickness varies from 0 to 48 m in the danish sector of the north sea (fig. 15). lithology. light grey to greenish grey, heavily bioturbated fig. 15. isochore map of the våle formation in the study area. the positions of the two danish reference wells, e-8 and siri-1, are indicated on the figure. 2047.5 2072.6 2186.5 2156.3 horda fm balder fm sele fm lista fm rind mb idun mb tyr mb vile mb ve mb bue mb bue mb våle fm chalk gp 2000 2100 2200 m siri-1 gr sonic neutron/densityfig. 14. siri-1, danish reference well for the våle and sele formations. black bars show cored sections. e-8 siri-1 0 10 20 30 40 thickness (m) våle formation 25 km 28 cm 2398 m 2399 m 2400 m cecilie-1b 2401 m 2402 m 0 10 20 30 40 50 60 70 80 90 100 cm underlying ekofisk formation is characterised by the ho of the dinoflagellate senoniasphaera inornata followed uphole by the ho of the planktonic foraminifer globoconusa daubjergensis. there is a hiatus at the contact between the ekofisk and våle formation in many sections and wells (clemmensen & thomsen 2005). the basal part of the våle formation, just above the top of the ekofisk formation, is marked by the downhole increase in calcareous foraminifer diversity and the ho of the planktonic foraminifers globanomalina cf. compressa and subbotina subdivision. the våle formation includes a sandstone unit (bor member, new) in the danish north sea sector. macroand ichnofossils. fragments of shelly macrofossils are present, but rare. the våle formation is heavily bioturbated. trace fossils in the formation include chondrites ispp., phycosiphon ispp., planolites ispp. and zoophycos ispp. thalassinoides ispp. burrows are only present locally. microfossils and palynomorphs. the uppermost part of the fig. 16. core photographs of the ekofisk–våle formation boundary in the cecilie-1b well. the shift from chalk to marlstones is gradational and placing the boundary can be difficult; it is positioned in the middle part of the core interval 2400.00–2401 m, at 2400.35 m (arrow), where light grey marls become dominant. depths are core depths. 29 trivialis. thus, the boundary between the two formations may in practice be located by reference to these three events (fig. 5a). the ho of the dinoflagellate alisocysta reticulata marks a level in the lower part of the våle formation. calcareous microfossil events near the top of the våle formation in the danish sector include the provisional ho of planktonic foraminifers. depositional environment. over most of the danish sector, the marlstones of the våle formation comprise hemipelagic deposits and deposits from dilute turbidity currents. the marlstones are probably largely of turbiditic origin. the foraminifer fauna of the våle formation is characterised by common calcareous taxa. taxa belonging to the neritic ‘midway-type’ fauna (berggren & aubert 1975) are especially common. the plankton/benthos ratio varies from approximately 1:1 in some areas to a total dominance of calcareous benthic foraminifers in other areas. the microfaunal composition indicates that the våle formation was deposited in an open marine, outer neritic environment that periodically reached upper bathyal depths. the bottom conditions were predominantly oxic with periods of dysoxia. the indications from the microfauna are supported by the trace fossil assemblage, which indicates water depths of at least 200 m combined with oxic to dysoxic bottom conditions. in the siri canyon, where thin turbidites are common in the våle formation, gravity flows played a major role during the deposition of the formation. age. selandian. correlation. the våle formation is equivalent to the lellinge greensand and the kerteminde marl onshore denmark and lithologically most closely resembles the latter. the oldest part of the kerteminde marl and the lellinge greensand are coeval, but the latter has a more restricted distribution (sjælland and storebælt regions only, fig. 1; thomsen 1994; clemmensen & thomsen 2005). alisocysta reticulata is consistently present in the lowest part of the kerteminde marl (clemmensen & thomsen 2005). the ho of a. reticulata is therefore an important intravåle as well as intra-kerteminde marl marker that may be used to correlate the two formations. the våle formation correlates with the marly facies of the maureen formation in the uk and norwegian sectors of the central and viking grabens (knox & holloway 1992). 2900 2800 2940.5 2903.0 2894.4 2913.3 2963.2 horda fm balder fm sele fm lista fm våle fm bor mb vile mb ve mb bue mb chalk gp m augusta-1 gr sonic neutron/densityfig. 17. augusta-1, type well for the bor, ve and bue members. black bars show cored sections. 30 våle fm chalk group vile mb bor mb lista fm augusta-1 clay si. vf. f. m. sand c. vc. c c c 45° 27° 2920 2930 2940 2950 2960 2920 2930 2940 2950 2960 log depth core depthgr sonic bor member new member history. the bor member encompasses sandstone bodies enveloped in the marlstones of the våle formation in the danish north sea sector. hardt et al. (1989) recognised a pure sandstone unit, the ty formation, located between the ekofisk and lista formations in the southern viking graben. the ty formation replaces the våle formation in its occurrence area and may be contemporaneous with the bor member, but it is not contiguous with it and it has a different source area. the presence of sandstone bodies in the våle formation in the danish sector was recognised by a stratigraphic working group at statoil norway in the mid-1990s and the sandstones were informally named the ‘borr member’. derivation of name. after bor (danish spelling), the father of odin. type well. danish sector well augusta-1, 2963.2–2940.5 m mdrt (figs 17, 18). reference well. danish sector well cecilie-1, 2319.6–2284.6 m mdrt (fig. 19). distribution and thickness. the bor member has been encountered at the mouth of the siri canyon as well as in fig. 18. core log of bor member sandstones in the augusta-1 well. legend in fig. 9. the interval around the våle–lista formation boundary marked by a grey bar in the core depth column is shown as core photographs in fig. 23. 31 the nearby wells tabita-1, augusta-1 and cleo-1 (fig. 20a; plate 1). it reaches a thickness of up to 23 m. lithology. the bor member consists of olive-green, partly calcite-cemented sandstones. the sandstones are very fine grained to fine grained and well sorted (fig. 18). rounded and translucent quartz grains dominate, but the content of glaucony grains is high (20–25%). mica and pyrite concretions are present in small amounts. angular chalk and claystone clasts occur locally. although composed exclusively of sandstone in the type well (fig. 18), the member may also include subordinate interbedded marlstones (e.g. cecilie-1, fig. 19). log characteristics. the bor member sandstones are best identified on the density log where they produce a blocky pattern with density values significantly lower than those of the marlstones beneath and above. the sandstones may also be identified from a combination of the density and neutron logs, as the presence of pure sandstone results in a ‘cross-over’ of the two log curves (figs 17, 19). on the gamma-ray log,thebormember ischaracterised by a blocky log signature with only small-scale increasing or decreasing trends and with values clearly higher than those of the subjacent, suprajacentand locally interbeddedmarlstones. boundaries. the boundaries with the marlstones of the våle formation, the chalks of the ekofisk formation and the mudstones of the lista formation are sharp and characterised by prominent shifts on the gamma-ray, sonic and density logs (figs 17–19, 21). depositional environment. the sandstones of the bor member were deposited from highly concentrated gravity flows at bathyal depths. age. selandian. correlation. the bor member is contemporaneous with parts of the kerteminde marl onshore denmark. the lellinge greensand, which appears between the top chalk surface and the kerteminde marl in some areas in eastern denmark, may also be broadly contemporaneous with the bor member, but differs from it lithologically in being predominantly a glaucony-rich calcilutite, rich in bryozoan fragments. the bor member may be compared with the ty formation (hardt et al. 1989) and with sandstones in the maureen formation (deegan & scull 1977) in the norwegian and uk sectors of the southern viking graben and the central graben. however, it is not contiguous with these units and it has a different source area. 2100 2200 2300 bue mb ve mb vile mb tyr mb bor mb horda fm balder fm sele fm lista fm våle fm chalk gp m cecilie-1 gr sonic neutron/density 2241.7 2276.6 2284.6 2319.6 fig. 19. cecilie-1, reference well for the bor and tyr members. black bar shows cored section. 32 fig. 20. location map showing the distribution of the rogaland group sandstones in the siri canyon (the outline of the canyon is indicated by grey shading, the grey shading inside the canyon indicates an area of positive relief within the canyon). a: bor and tyr members. b: idun member. c: rind and kolga members. cleo-1 francisca-1 frida-1 d-1 nolde-1 siri-3siri-2 connie-1 siri-1 augusta-1 elna-1 tabita-1 amalie-1 cecilie-1 10 km augusta-1 tabita-1 nini-3 amalie-1 cleo-1 elna-1 frida-1 d-1 nolde-1 cecilie-1 siri-3 sandra-1 sandra-1 siri-2 siri-1 connie-1 10 km augusta-1 tabita-1 amalie-1 cleo-1 elna-1 frida-1 d-1 nolde-1 cecilie-1 siri-3 sandra-1 siri-2 siri-1 connie-1 10 km tyr mb bor mb both mbs idun mb rind mb kolga mb both mbs sir i c an yo n sir i c an yo n sir i c an yo n a b c nini-3 francisca-1 francisca-1 nini-3 lista formation history. deegan & scull (1977) established the lista formation for the widespread, non-laminated mudstones that overlie the marls of the unit equivalent to the maureen formation (våle formation). kristoffersen & bang (1982) established the non-calcareous clay and shale unit cen-1 between the top of their north sea marl (våle formation) and the base of the beds with volcanic tuff. they noted that the cen-1 unit corresponds to the lista formation. for reasons of seniority and the informal nature of the cen units, we maintain the name lista formation for this stratigraphic unit. type well. norwegian sector well 2/7–1, 2917.5–2872.5 m mdkb. danish reference wells. e-8, 2057.0–2027.6 m mdkb (fig. 13); cleo-1, 2812.0–2765.5 m mdkb (fig. 21; plate 1). distribution and thickness. the lista formation is present throughout the north sea basin, except in a few areas where it has been removed by erosion. in the danish sector, its thickness varies from 0 to 108 m (fig. 22). lithology. the formation is characterised by dark coloured, predominantly greyish, greenish or brownish, non-laminated to faintly laminated, non-calcareous mudstones. the lista formation is predominantly non-tuffaceous but becomes tuffaceous towards its top. in the siri canyon, glaucony-rich, massive sandstone layers and injected sandstone bodies occur in the lista formation. log characteristics. although fluctuating, both the gamma-ray and sonic log readings in the lista formation have higher mean values than those of the underlying våle formation and lower mean values than those of the overlying sele formation. in wells where mudstone facies dominate in the lista formation, the gamma-ray and sonic log patterns can be subdivided into three. the tripartite log pattern reflects the succession of three different mudstone units, established as new members herein (see below). boundaries. in most wells where the transition has been cored, the boundary is sharp between the light-coloured marlstones of the våle formation and the dark-coloured, non-calcareous mudstones of the lower lista formation (vile member, see below; fig. 23). on the gamma-ray log, the boundary is picked at an abrupt upward shift to higher values than in the underlying våle formation. this level can typically be identifiedon the sonic log at a velocity 33 minimum. above this minimum, the sonic readings increase slightly upwards. the lista formation is overlain by the sele formation. the base of the sele formation was defined by deegan & scull (1977 p. 34) at the contact between “non-laminated, non-tuffaceous shales” (lista formation) and “laminated tuffaceous shales” (sele formation). this boundary definition was followed by mudge & copestake (1992a, b). on the other hand, knox & holloway (1992 p. 46) followed o’connor &walker (1993) and placed the boundary somewhat lower, at the contact between “grey-green and green-grey, blocky, bioturbated claystones” of the lista formation and “dark grey fissile mudstones” of the sele formation. the boundary concept of knox & holloway implies that the “non-laminated, non-tuffaceous shales” of deegan & scull are incorporated in the sele formation where these, together with overlying laminated indisputable sele mudstones, constitute the basal sele unit s1a (knox & holloway 1992). in the present paper, the boundary concept of deegan & scull (1977) is followed, and the lower part of the unit of “non-laminated, nontuffaceous shales” (the “dark grey, fissile mudstones” of knox & holloway) that overlies the grey-green mudstones is retained in the lista formation as its topmost unit. this unit is formalised as a new member of lista formation herein (bue member, see below). subdivision. the lista formation is subdivided into six new members. three of these, the vile, ve and bue members, are mudstone units that have widespread distribution in the north sea basin and can be correlated with danish onshore units. in the siri canyon, fine-grained fig. 22. isochore map of the lista formation in the study area. the positions of the two danish reference wells, cleo-1 and e-8, are indicated. 2700 2800 bue mb ve mb vile mb bor mb horda fm balder fm sele fm lista fm våle fm chalk gp m cleo-1 2765.5 2812.0 2792.4 2777.6 gr sonic densityfig. 21. cleo-1, danish reference well for the lista formation and the vile, ve and bue members. lista formation cleo-1 e-8 20 40 60 80 100 thickness (m) 25 km 34 2918 m 2919 m augusta-1 2920 m 2921 m 0 10 20 30 40 50 60 70 80 90 100 cm sandstone bodies occur within each of the three mudstone units. these sandstone-dominated units are proposed here as three new members: the tyr member for the sandstones in the vile member, the idun member for the sandstones in the ve member, and the rind member for the sandstones in the bue member. knox & holloway (1992) recognised a threefold subdivision of the lista formation exclusively based on biostratigraphy. their l1 and l2 units are separated by the ho of the dinoflagellate palaeoperidinium pyrophorum, and the l2 and l3 units are separated by the ho of areoligera gippingense. it is noticeable that these two bioevents occur close to the boundaries between the three mudstone members of the lista formation proposed herein on the basis of lithology. macroand ichnofossils. macrofossils have not been reported fromthelistaformation;theformationismoderatelytoheavily bioturbated(for ichnotaxa, see individualmembersbelow). microfossils and palynomorphs. the lista formation differs from the underlying våle formation by lacking common planktonic foraminifers and from the overlying sele formation by having an impoverished agglutinated benthic foraminifer assemblage. the lista formation contains a characteristic sequence of palynomorph datums that can aid separation of its members. these datums are treated under the individual lista members below. depositional environment. the lista formation consists predominantly of hemipelagic mudstones and was probably deposited from very dilute turbidity currents and from suspension. the composition of the microfaunal assemblage indicates a relatively open marine depositional setting in upper to possibly middle bathyal depths with oxic to dysoxic bottom conditions. this is based on the presence of an impoverished agglutinated foraminifer assemblage dominated by tubular suspension feeders (especially rhabdammina spp.) together with epifaunal and infaunal detritivores (e.g. haplophragmoides spp. and spiroplectammina spectabilis). the relative abundance of tubular suspension feeders is higher in wells in the siri canyon than in wells outside the canyon. this probably indicates slightly deeper water within the canyon area during deposition of the lista formation. age. selandian–thanetian (upper paleocene). the selandian–thanetian boundary may be placed in the middle part of the lista formation (in the lower part of the ve member, see below), at the ho of the dinoflagellate palaeoperidinium pyrophorum. correlation. the lista formation corresponds to the following succession of upper paleocene units from onshore fig. 23. core photographs of the våle–lista formation boundary interval in the augusta-1 well. the boundary is at 2919.46 m where dark grey non-calcareous mudstones of the vile member overlie greenish grey marls of the våle formation. in most cores the boundary is sharp, as illustrated here. depths are core depths. stratigraphic position of the figured interval is shown in fig. 18. 35 fig. 25. correlation diagram showing possible diachronism of the våle–vile boundary in an east–west transect extending into the norwegian sector of the north sea (1/3-1 and 2/7-1 are norwegian sector wells). the distribution of biostratigraphic events shows that the våle–vile boundary youngs in a westerly direction (ho, highest occurrence). alternatively, the event distribution could be explained as a result of reworking of older strata into the lista formation in the danish sector. the figure also shows an example of a well (2/7-1) with a relatively large separation between the base of the sele formation and the lowest and most conspicuous gamma-ray peak in the formation (see text for further explanation). æbelø fm holmehus fm østerrende clay ølst fm kerteminde marl danian limestone 1/3-1 gr sonic gr sonic gr sonic gr sonic 1/3-1 2/7-1 mona-1 e-8 viborg-1 viborg-1 sele fm lista fm bue mb ve mb vile mb våle fm chalk group ho abundant palaeoperidinium pyrophorum ho isabelidinium? viborgense ho diverse calcareous benthic foraminifers ho (provisional) planktonic foraminifers 2050 2075 450 400 3050 3000 2900 2950 3150 3200 3250 kolga mb s3 s2b s2a s1b bue mb ve mb vile mb horda fm balder fm sele fm lista fm våle fm chalk gp 2000 2100 m siri-3 1998.8 2016.8 2036.1 2066.4 2072.8 2102.6 gr sonic densityfig. 24. siri-3, type well for the vile and kolga members, and danish reference well for the balder formation. the figure also shows the subdivision of the sele formation used by knox & holloway (1992); in this well, the bue member is equivalent to the s1a subunit of these authors. black bar shows cored section. 36 denmark: æbelø formation (informal mudstone unit described by bøggild 1918 and heilmann-clausen 1995), holmehus formation (heilmann-clausen et al. 1985) and østerrende clay (informal mudstone unit described by nielsen et al. 1986 and heilmann-clausen 1995). vile member new member history. the vile member comprises the widespread, dark olive-grey to dark grey, non-calcareous, fissile mudstones that constitute the lower part of the lista formation. the unit was recognised by a stratigraphic working group at statoil norway in the mid-1990s and informally named the ‘vile formation’. derivation of name. after vile, the brother of odin. type well. danish sector well siri-3, 2102.6–2072.8 m mbrt (fig. 24; plate 4). reference wells. danish sector wells e-8, 2057.0–2044.0 m mdkb (fig. 13); cleo-1, 2812.0–2792.4 m mdkb (fig. 21; plate 1). distribution and thickness. the vile member has been recognised in a large number of north sea wells, and the unit probably has a basinwide distribution. however, it is apparently lacking in the siri canyon wells connie-1 and siri-2 (figs 29, 31), probably due to erosion. its thickness varies between 0 and 30 m over most of the danish sector. it greatest thickness is reached in the siri canyon. lithology. the member consists of dark olive-grey to dark grey, non-calcareous, swelling, smectitic, fissile mudstones (fig. 23). thin silicified layers occur in the member. calcite is common and occurs as small nodules and larger concretions. in the siri canyon, small pyrite concretions and less than 1 cm thick, silty, very fine-grained glaucony-rich sandor siltstone laminae are locally present in the vile member. the laminae are parallel to the bedding of the mudstones; they have sharp bases and are normally grad1803.7 1763.9 1700.4 1717.2 horda fm lark fm balder fm sele fm kolga mb bue mb vile mb idun mb tyr mb lista fm chalk gp 1600 1700 1800 m nini-3 gr sonic neutron/density s3 s2b s2a s1b fig. 26. nini-3, type well for the tyr member and reference well for the kolga member. the figure shows the subdivision of the sele formation used by knox & holloway (1992); in this well, the bue member is equivalent to the s1a subunit of these authors. black bar shows cored section. 37 ed. thin concordant or discordant, postdepositional sandstone intrusions are locally present. in the lower part of the vile member, the intrusions are only a few millimetres thick, but they often increase in number and thickness towards the top of the member (fig. 11). log characteristics. in most wells there is a gradual increase in gamma-ray response up through the vile member, accompanied by a slight decrease in sonic readings. boundaries. the lower boundary of the vile member is that of the lista formation. the upper boundary is defined by the base of the ve member; boundaries with the sandstone-dominated tyr member are described under that member. macroand ichnofossils. the vile member is moderately to intensely bioturbated. ichnogenera in the member include chondrites ispp., phycosiphon ispp., planolites ispp. and zoophycos ispp. microfossils and palynomorphs. the vile member is characterised by a general decrease in the diversity of benthic foraminifers and radiolaria from its base to its top. in the danish sector, the transition from the underlying våle formation to the vile member is marked by the provisional ho of planktonic foraminifers. a conspicuous drop in diversity of benthic foraminifers takes place in the middle of the vile member. the ho of the dinoflagellate isabelidinium? viborgense is an important intra-vile marker located in the upper part of the member. above it, a sudden decrease in the abundance of radiolaria further charlista fm chalk group nini-3 clay si. vf. f. m. sand c. vc. p p p 1800 1790 1780 1770 1760 1800 1790 1780 1770 1760 log depth core depthgr sonic vile mb tyr mb fig. 27. core log of the tyr member in the nini-3 well. for legend, see fig. 9. 38 2908 m2907 m 2909 m augusta-1 2910 m 2911 m 2912 m 0 10 20 30 40 50 60 70 80 90 100 cm acterises a level within the uppermost part of the vile member. the transition from the vile member to the overlying ve member is marked by a conspicuous drop in the abundance of the dinoflagellate palaeoperidinium pyrophorum(whichhas itshoataslightlyhigherstratigraphic level, within the lower part of the ve member, see below). depositional environment. the mudstones of the vile member are hemipelagic deposits, whereas the thin sandstone and siltstone laminae are interpreted as the deposits of low-density turbidity currents. the presence of zoophycos ispp. suggests depositional water depths of at least 200 m (bottjer & droser 1992). age. selandian. correlation.the vile member corresponds to theæbelø formation, onshore denmark (informal mudstone unit describedbybøggild1918andheilmann-clausen1995). it corresponds to the lista l1 subunit of knox & holloway (1992). fig. 28. core photographs of mudstones of the vile and ve members in the augusta-1 well. depths are core depths. the boundary between the two members is placed where greenish and reddish grey mudstones become dominant, at 2910.4 m (arrow). this depth corresponds to log depth 2913.3 m on fig. 17. 39 biostratigraphic correlation with well sections in the norwegian north sea sector may indicate that the lower boundaryof thevilemember isdiachronous (fig. 25). in the type well for the lista formation (norwegian well 2/7-1; fig. 25), its base (i.e. its contact with the marlstones of the underlying våle formation) is above the ho of a diverse calcareous benthic foraminifer assemblage. in wells in the danish sector, this event occurs within the vile member. similarly, the provisional ho of planktonic foraminifers, an event that is close to the boundary between the våle and lista formations in denmark, is found well within the våle formation in well 2/7-1. in the type well for the våle formation (norwegian well 1/3-1; fig. 25), the ho of i.? viborgense coincides with the våle–lista boundary (i. prince, unpublished biostratigraphic data). this event occurs above the våle formation in the danish sector, in the middle to upper part of the vile member. in the danish onshore well viborg-1, the latter event and the ho of the diverse calcareous benthic foraminifer assemblage occur above the kerteminde marl in the upper part of the æbelø formation, a correlative of the vile member (fig. 25; heilmann-clausen 1985). the distribution pattern of the biostratigraphic events indicates that sedimentation of marls continued in the norwegian sector some time after marl sedimentation was replaced by sedimentation of non-calcareous mudstones in the danish sector. alternatively, calcareous foraminifer assemblages and associated lithologies have been reworked into higher levels of the våle formation or even into the lista formation in the danish sector. tyr member new member history. the tyr member consists of glaucony-rich, sandstone-dominated deposits that are laterally equivalent to, and commonly underlain and overlain by, mudstones of the vile member. these sandstones were previously recognised by a stratigraphic working group at statoil norway in the mid-1990s and informally referred to the ty formation of hardt et al. (1989). derivation of name. after tyr, the son of odin. type well. danish sector well nini-3, 1803.7–1763.9 m mdrt (figs 26, 27; plate 4). reference well. danish sector well cecilie-1, 2276.6–2241.7 m mdrt (figs 8, 19). distribution and thickness. the tyr member has only been encountered in the siri canyon and it may be restricted to thatarea. it reaches a thickness of up to 40 m (fig. 20a). lithology. the tyr member is characterised by thick beds of olive-green to greenish grey, very fine-grained to finegrained and well-sorted sandstone (fig. 27). rounded and translucent quartz grains dominate, but the content of glaucony grains is high (15–20%), hence the greenish colour of the sandstones. mica and small pyrite concretions are present in small amounts throughout the member. angular chalk and claystone clasts occur locally in the sandstones. the sandstones are partly calcite-cemented. intrusive sandstones are common, particularly towards the top of the member where they may be several metres thick (fig. 8). subordinate interbedded dark grey noncalcareous mudstones resemble those of the laterally equivalent vile member. log characteristics. the tyr member is best identified on the density log where the sandstones are characterised by a conspicuously lower density than the associated mudstones. the sandstones may also be identified from a combination of the density and neutron logs, as the presence of pure sandstone results in a ‘cross-over’ of the two log curves (figs 19, 26). the gamma-ray response resembles that of the underlying våle formation, but is slightly lower than the response of the vile member (figs 8, 26). this log pattern makes it feasible to differentiate even minor sandunits from mudstone beds in the tyr member. thicker sand units may show decreasingor increasingupwards gamma-ray values. these trends do not seem to berelatedtograin-sizevariations, judgingfromcore studies. boundaries. the boundaries to the mudstones of the vile member, the marlstones of the våle formation and the chalks of the ekofisk formation are sharp and characterised by prominent shifts in gamma, sonic and density log readings (figs 19, 26, 27). in some wells, the tyr member overlies the våle formation or the ekofisk formation with an erosional contact (e.g. nini-3; figs 26, 27). depositional environment. although the sandstones of the tyr member were deposited from highly concentrated gravity flows, their present appearance is dominated by theeffectsofpostdepositional liquefaction and fluidisation. age. selandian. correlation. the tyr member is contemporaneous with parts of the lithologically dissimilar æbelø formation in 40 onshore denmark and with the lower part of the heimdal formation of deegan & scull (1977) as well as the andrew sandstone and the mey sandstone member of knox & holloway (1992) in the norwegian and uk sectors of the southern viking graben. however, it is not contiguous with the latter three sandstone units and has a different source area. ve member new member history. the ve member consists of variegated mudstones that have previously been recognised from north sea wells as the holmehus formation by heilmann-clausen et al. (1985), who gave no further details, and by danielsen & thomsen (1997), who indicated its presence in several wells. the unit was also recognised by a stratigraphic working group at statoil norway in the mid-1990s and informally named the ‘ve formation’. derivation of name. after ve, the brother of odin. type well. danish sector well augusta-1, 2913.3–2903.0 m mdrt (fig. 17). reference wells. danish sector wells e-8, 2044.0–2030.3 m mdkb (fig. 13); cleo-1, 2792.4–2777.6 m mdkb (fig. 21; plate 1). distribution and thickness. the sediments of the ve member have been recognised from a large number of north sea wells, and the unit probably has an almost basinwide distribution. its thickness varies from 0 to 21 m in the danish sector. lithology. the ve member consists of mottled green, bluish green, reddish brown and brown mudstones (fig. 28). mottled, purple coloured intervals are also present locally. the middle part of the member is often characterised by a thick dark reddish brown to chocolate brown interval. thevemembermudstones are non-calcareous and richin smectite. pyrite and carbonate concretions occur throughout the member. a weak biogenic lamination is sometimes observed in cores. only very little organic material is present in the member. in the siri canyon, thin intrusive sandstones are common in the ve member. log characteristics. in general, the gamma-ray log shows a decreasing-upwards trend through the ve member, as opposed to the increasing trend through the underlying vile member. in the uppermost part of the ve member, the gamma-ray response increases over a short interval before reaching the base of the overlying bue member. the sonic log pattern throughout the ve member is smooth and relatively stable compared with the sonic pattern of the vile member. it also differs from the latter in having an increasing-upwards trend. the log pattern of the ve member differs from that of the overlying bue member in having a lower gamma-ray response level. boundaries. the lower boundary, with the vile member, is placed at the first appearance of greenish, bluish or reddish brown mudstones above the dark olive-grey to dark grey mudstones of the vile member. the colour change from vile to ve mudstones is often gradational and the boundary may be difficult to define precisely (fig. 28), especially when only cuttings samples are available. however, in the colour transition interval, a gamma-ray spike separates an interval with an increasing-upwards gammaray trend below from an interval with a decreasing gamma-ray trend above (compare figs 17 and 28). in the absence of a clear indication of the boundary level from sediment colour change, the gamma spike at the shift from increasing to decreasing gamma-ray values may be used as a marker for the boundary. the vile member is absent from the connie-1 well, and in this well the lower contact of the ve member is with the marlstones of the våle formation (fig. 31). the upper boundary is at the base of the bue member (see below). boundaries with the idun member are described under that member. macroand ichnofossils. the mudstones of the ve member are normally heavily bioturbated. the most common trace fossils are phycosiphon ispp. and zoophycos ispp. chondrites ispp. and planolites ispp. are present, but rare. microfossils and palynomorphs. in the danish sector of the north sea, the ho of abundant palaeoperidinium pyrophorum is located at or close to the vile–ve boundary. the hos of p. pyrophorum and palaeocystodinium australinum are in the lower part of the ve member. in general, the dinoflagellate assemblage from the upper part of the ve member is sparse and is characterised by specimens of areoligera gippingensis. an acme of the latter species marks a level in the upper part of the ve member. the highest in situ occurrence of the dinoflagellate alisocysta margarita is located close to the top of the ve member. depositional environment. deposition of the mudstones of the ve member was controlled by hemipelagic sedimentation and sedimentation from dilute turbidites. the 41 idun member new member history. the idun member consists of sandstone-dominated deposits that are laterally equivalent to, and commonly underlain by, mudstones of the ve member. this sandstone unit was previously recognised by a stratigraphic working group at statoil norway in the mid-1990s and was informally referred to the heimdal formation of deegan & scull (1977). derivation of name. after idun, the goddess of youth. type well. danish sector well connie-1, 2368.3–2332.0 m mdrt (figs 29, 30). 2200 2300 2368.3 2332.0 2329.6 2292.2 idun mb rind mb bue mb bue mb ve mb horda fm balder fm sele fm lista fm våle fm chalk gp m connie-1 gr sonic neutron/densityfig. 29. connie-1, type well for the idun and rind members. in this well, the rind member may be divided into three major sandstone intervals. the idun member consists of two thick sandstone intervals, separated by a thick mudstone unit. black bars show cored sections. occurrence of the trace fossil zoophycos ispp. indicates a water depth of at least 200 m (bottjer & droser 1992). the overall high degree of bioturbation, the lack of organic material and the greenish, bluish and reddish brown colours together suggest oxygenated bottom conditions. age.selandian–thanetian.theselandian–thanetianboundaryisplacedatthehoofthedinoflagellatepalaeoperidinium pyrophorum, in the lower part of theve member. correlation. the ve member correlates with the holmehus formation (heilmann-clausen et al. 1985) onshore denmark and is lithologically indistinguishable from that formation. 42 lista fm våle fm connie-1 clay si. vf. f. m. sand c. vc. 45o s v g g g pp ppp 45o 2290 2300 2310 2320 2340 2350 2360 2300 2290 2310 2320 2330 2340 2350 2360 2370 log depth core depthgr sonic p p p p p p p p p p p p change of core depth scale bue mb idun mb bue mb rind mb ve mb fig. 30. core log of the idun and rind members in the connie-1 well. for legend, see fig. 9. minor mudstone beds separate the three major sandstone intervals of the rind member. minor mudstone layers are also intercalated with the idun member sandstone intervals. the core depth scale of the lower core is offset by c. 1.6 m relative to the scale of the upper core in the figure. for reasons of consistency with core data from this well, the original (albeit erroneous) core depths of the lower core are maintained in the figure. this does not affect the depths of the top and base of the idun member given in the text, as these are based on log depths. 43 reference well. danish sector well siri-2, 2205.5–2127.0 m mdrt (fig. 31). distribution and thickness. the idun member is only known from the siri canyon, and it may be restricted to that area (fig. 20b). it reaches a thickness of up to 179 m in the siri-2 well. lithology. the idun member is dominated by very finegrained to fine-grained, well-sorted sandstones (fig. 30). rounded and translucent quartz grains dominate, but the content of glaucony grains is high (15–25%). the sandstones are olive green to greenish grey due to the high content of glaucony. mica and small pyrite concretions are present in small amounts. angular chalk and claystone clasts occur locally. intrusive sandstones are also represented (figs 11, 12, 30). the sandstones and adjacent mudstones are partly calcite-cemented. in the nini and siri wells, the sandstones occur in thick amalgamated successions with only rare, thin mudstone interbeds; the latter are lithologically comparable to the laterally equivalent ve member mudstones (see above). in the connie-1 well, however, the sand-rich succession is interrupted by a discrete 6 m thick mudstone unit (fig. 30). log characteristics. the sandstone-dominated idun member is best identified on the density log where it is characterised by a conspicuously lower density than the underlying and overlying mudstones (figs 29, 31). the sandstone component may also be identified from a combination of the density and neutron logs, as the presence of pure sandstones results in a ‘cross-over’ of the two log curves (figs 29, 31). the idun member is characterised by a blocky, decreasing-upwards gamma-ray and density log pattern. intervals with an overall constant gamma-ray pattern may be characterised by many small-scale increasingor decreasing-upwards gamma-ray cycles. boundaries. in sections where the sandstones of the idun member are enveloped by mudstones of the ve member, the boundaries are sharp and characterised by prominent shifts on the gamma-ray, sonic and density logs (figs 29, 30). where the ve member is absent, comparable, sharp boundaries are observed with the mudstones of the vile member beneath and the bue member above (figs 11, 26, 30). in the siri-2 well, the lower lista formation is absent and an erosive unconformity separates the idun member sandstones from the marlstones of the våle formation (fig. 31). depositional environment. although the sandstones of the idun member were deposited from highly concentrated gravity flows, their present appearance largely records postdepositional liquefaction and fluidisation processes. age. selandian–thanetian. bue mb bue mb rind mb idun mb balder fm horda fm sele fm lista fm chalk gp våle fm 2100 2200 m siri-2 gr sonic neutron/density 2127.0 2205.5 fig. 31. siri-2, reference well for the idun member. black bar shows cored section. 44 with these sandstone units, however, and has a different source area. bue member new member history. the bue member encompasses the light to dark grey and greyish black mudstones that occur between the top of the ve member and the base of the sele formation. these mudstones have not previously been recognised as a separate unit in the danish sector. derivation of name. after bue, the son of odin and rind. typewell. danishsectorwellaugusta-1,2903.0–2894.4m mdrt (fig. 17). referencewells.danishsector wells e-8, 2030.3–2027.6 m below mdkb (fig. 13); cleo-1, 2777.6–2765.5 m mdkb (fig. 21; plate 1). distribution and thickness. the sediments of the bue member have been recognised from a largenumberof north sea wells, and the unit probably has a basinwide distribution. its thickness varies from 0 to 18 m in the danish sector. lithology. the bue member consists of light to dark grey and greyish black mudstones. the mudstones are generally rich in smectite. in the siri canyon, the upper part of the member sometimes contains laminae of very finegrained to fine-grained sandstones or siltstones, mimicking the laminated mudstones of the overlying sele formation (figs 32, 39). the laminae are less than 2 cm thick, have sharp bases, are normally graded and show parallel lamination. concordant or discordant sandstone intrusions are locally present in the member (fig. 8). small calcite and siderite concretions are occasionally present. moderately to intensely bioturbated intervals are interbedded with non-bioturbated intervals. tuff layers may be present in the member. log characteristics. the gamma-ray response of the bue member is generally higher than that of the underlying ve member, but lower than that of the overlying sele formation. in some siri canyon wells, minor coarseningupwards cycles are indicated by the gamma-ray log of the bue member. boundaries. the transition from typical lithologies of the ve member to those of the bue member is often gradafig. 32. core photographs showing the bue member mudstones with numerous sandstone laminae in the augusta-1 well. depths are core depths. 0 10 20 30 40 50 60 70 80 90 100 cm 2891 m 2892 m augusta-1 2893 m correlation. the idun member is contemporaneous with parts of the lithologically dissimilar holmehus formation onshore denmark and with parts of the heimdal formation of hardt et al. (1989) and the lower balmoral sandstone and tuffite of the mey sandstone member of knox & holloway (1992). the idun member is not contiguous 45 rind member new member history. the rind member consists of sandstone-dominated deposits that are laterally equivalent to, and commonly underlain and overlain by, mudstones of the bue member. sandstone bodies at this stratigraphic level were previously recognised by a stratigraphic working group at statoil norway in the mid-1990s and were informally referredtotheheimdalformationofdeegan&scull (1977). derivation of name. after the giantess rind. type well. danish sector well connie-1, 2329.6–2292.2 m mdrt (figs 29, 30). reference well. danish sector well sandra-1, 2066.3– 2004.8 m mdrt (fig. 33). distributionand thickness. the rind member has only been encountered in the siri canyon, and it may be restricted to that area where it reaches a thickness of 62 m (fig. 20c). lithology. the rind member consists of very fine-grained, well-sorted sandstones interbedded with thin mudstone beds that typically form less than 15% of the member (fig. 30). rounded and translucent quartz grains dominate in the sandstones, but the content of glaucony grains in the very fine-grained to fine-grained size fraction is high (15–25%). the sandstones are olive green to greenish grey due to the high content of glaucony. mica and small pyrite concretions are present in small amounts. angular chalk and claystone clasts occur locally in the sandstones, which are partly calcite-cemented. the interbedded mudstones are lithologically comparable to the bue member mudstones (see above). log characteristics. the rind member is best recognised on the density log where it shows either a blocky or a serrate pattern created by the alternation of sandstone beds or amalgamated units (low density) with thin mudstone beds (high density; figs 29, 33). the sandstones may also be identified from a combination of the density and neutron logs, since the presence of pure sandstones results in a ‘cross-over’ of the two log curves (figs 29, 33). the gamma-ray log shows a low-amplitude serrate pattern. this pattern does not reflect alternating sand or mudstones, judging from core inspection. boundaries. the boundary between the sandstones of the rind member and the mudstones of the bue member is tional and the boundary may therefore be difficult to position precisely. it is placed where mottled green, bluish green, reddish brown and brown mudstones pass upwards into grey mudstones with sandstone and siltstone laminae. on the petrophysical logs, this transition is reflected by a shift from decreasingto increasing-upwards gamma-ray values or at an abrupt increase in the gamma-ray response. the upper boundary of the bue member is at the base of the sele formation. macroand ichnofossils. trace fossils recognised in the bue member are phycosiphon ispp., planolites ispp., thalassinoides ispp. and rare zoophycos ispp. microfossils and palynomorphs. the ve–bue boundary is bracketed by the stratigraphic succession of the ho of in situ alisocysta margarita (occurring in the upper ve member) followed by the ho of an impoverished assemblage of benthic agglutinated foraminifers (in the lower part of the bue member). the upper part of the bue member is characterised by common spores and pollen, in particular bisaccate pollen and inaperturopollenites spp. the bue– sele boundary is marked by the base of an acme of the dinoflagellate genus apectodinium and the lo of apectodinium augustum. depositional environment. the normally graded sandstone to siltstone laminae in the upper part of the member indicate that deposition of the bue member took place from dilute, low-density, turbidity currents in a generally sediment-starved environment at this level. the minor coarsening-upwards cycles observed on petrophysical logs from some siri canyon wells probably indicate either small distal lobes of deep-water channel-sandstones or levee deposits. age. thanetian. correlation. the bue member corresponds to the østerrende clay (informal mudstone unit described by nielsen et al. 1986 and heilmann-clausen 1995) onshore denmark. the level here defined as the boundary between the ve and thebuememberswascorrelatedbyknox (1997 fig. 3; the lista–sele boundary of this worker) with the boundary between the holmehus formation and the østerrende clay (as ‘grey clay’) onshore denmark. the bue member further correlates with the lower part of the s1a subunit of the sele formation established by knox & holloway (1992; see correlation section under the lista formation for further details). 46 sharp and characterised by prominent shifts on both the sonic and density logs (figs 29, 33). it is often difficult to identify the boundaries on the gamma-ray log alone. depositional environment. although the sandstones of the rind member were deposited from highly concentrated gravity flows, their present appearance largely records postdepositional liquefaction and fluidisation processes. age. thanetian. correlation. the rind member may be contemporaneous with parts of the lithologically dissimilar østerrende clay encountered in the storebælt region (fig. 1), with sandstones in the higher parts of the heimdal formation (hardt et al. 1989) and with the upper balmoral sandstone of the mey sandstone member of knox & holloway(1992).however, therindmember is not contiguous with those sandstone units and has a different source area. sele formation history. the sele formation was established by deegan & scull (1977) for the dark grey to greenish grey, laminated and carbonaceous, tuffaceous, montmorillonite-rich shales and siltstones that overlie the non-laminated and nontuffaceous shales of the lista formation in some areas, or arenaceous sediments belonging to a variety of different units in other areas. the original definition of the sele boundary is followed herein. this implies that the base of the sele formation is located at the base of the “laminated tuffaceous shales” that overlie the “non-laminated, nontuffaceous shales” of the lista formation (deegan & scull 1977; see boundaries section under the lista formation for further details). sandstones occur in the sele formation in the danish sector; these are established as a new member, the kolga member. typewell.british sectorwell21/10-1,2131–2100m mdkb. bue mb rind mb vile mb ve mb balder fm horda fm lark fm sele fm lista fm våle fm 1800 1900 2000 2004.8 2066.3 2100 m sandra-1 gr sonic neutron/density fig. 33. sandra-1, reference well for the rind member. black bar shows cored section. 47 danish reference wells. siri-1, 2072.6–2047.5 m mdkb (fig. 14; plates 1, 4); tabita-1, 2958.8–2941.4 m mdkb (fig. 34; plate 2). distribution and thickness. the sele formation is recognised from a large number of north sea wells and it has a basinwide distribution. in the danish sector, the thickness varies from 5 to 54 m (fig. 35). lithology. the sele formation consists of medium to dark grey, brownish or black laminated mudstones. thin tuff layers occur in the upper part of the formation. it contains three or more well-laminated intervals where dark mudstone beds alternate with lighter coloured mudstone beds. the well-laminated intervals are enriched in organic material resulting in a high gamma-ray response, primarily due to increased uranium content. the most organic-rich, and often darkest, most well-laminated interval is found in the basal part of the formation (fig. 36). the mudstones of the sele formation show an overall upward increase in the silt fraction. in the upper half of the formation, the mudstones may be interbedded with thin, very fine-grained sandstone laminae and thin sandstone beds (fig. 37). the sandstone beds are up to 12 cm thick, normally graded and display parallel lamination. locally, and dominantly in the upper part of the formation, graded tuff laminae less than 1 cm thick are present. in cores, the tuff laminae have a light purple colour. small calcite concretions are present, but rare. in the siri canyon, the sele formation is interbedded with sandstones or it grades upwards into a succession of thinly interbedded sandstones and mudstones. thin sandstone intrusions occur, but only in the lower part of the formation. log characteristics. the sele formation is characterised by high gamma-ray readings throughout, with a number of gamma-ray peaks. on the gamma-ray log, the base of the sele formation is generally marked by a conspicuous upward shift to consistently higher gamma-ray readings than those of the underlying bue member (figs 38, 39). in most wells, a pronounced gamma-ray peak follows a short distance above the base of the sele formation (e.g. augusta-1 and e-8; fig. 38). in wells to the north and west of the danish sector, the stratigraphic distance between the shift to higher gamma-ray readings at the base of the sele formation and the gamma-ray peak is considerably greater (e.g. in the norwegian well 2/7-1; fig. 25). in some wells in the danish sector (and in most siri canyon wells), the basal high gamma-ray interval is missing and the base of the sele formation is marked by the pronounced gamma-ray peak (e.g. cleo-1 and nini-3; figs 38, 39). boundaries. the lower boundary is characterised by a change from the light to dark grey and greyish black mudstones with thin sandstone laminae of the bue member (lista formation), to dark grey to black well-laminated mudstones without sandstone laminae of the sele formation (fig. 36). the upper boundary is at the base of the balder formation. subdivision. knox & holloway (1992) suggested an informal threefold subdivision of the sele formation based fig. 35. isochore map of the sele formation in the study area. the positions of the two danish reference wells, siri-1 and tabita-1, are indicated in the figure. siri-1 tabita-1 10 20 30 40 50 thickness (m) sele formation 25 km fig. 34. tabita-1, danish reference well for the sele formation. horda fm balder fm sele fm lista fm vile mb ve mb bue mb våle fm chalk gp 2900 3000 m tabita-1 2941.4 2958.8 gr sonic 48 in figs 24, 26 and 38), whereas other lithological changes are more subtle. the subdivision is outlined below. unit s1 this unit comprises a lower subdivision s1a and an overlying subdivision s1b. subdivision s1a is identical to the bue member of the lista formation and the stratigraphic interval from the base of the sele formation (sensu deegan & scull 1977) up to the base of the lowermost conspicuous gamma-ray peak within that formation (fig. 38). that peak is associated with relatively low sonic values. subdivision s1b has its base at the gamma-ray peak and its top at the base of the next gamma-ray peak. the gamma-ray response decreases up through s1b. unit s1 consists of brownish grey to dark grey and black, well-laminated mudstones. unit s2 the base of the unit is at the base of the second gamma-ray peak. this peak can be differentiated from the gamma-ray peak at the base of s1b by its association with high sonic values. unit s2 can be divided into a lower subdivision (s2a) characterised by a relatively high gamma-ray response level and an overlying subdivision (s2b) with a lower gamma-ray response (fig. 38). the two subdivisions are separated by a gammaray low. unit s2 consists of light grey to brownish dark grey, laminated to well-laminated mudstones. both the lamination and the colour of the two subdivisions are very similar in the cores encountered in this study and it is almost impossible to distinguish the two subdivisions on lithology alone. knox & holloway (1992) observed tuff layers in the basal part of unit s2b in the british wells; tuffs were not observed in the danish wells.. unit s3 this unit is characterised by high and increasingupwards gamma-ray values. the base of the unit is defined by a sharp increase in gamma-ray readings (fig. 38). there is no significant colour difference between the mudstones of units s2 and s3 in the danish wells, but lamination seems to be better developed in unit s3 than in unit s2. tuff layers were observed in unit s3 in the british wells studied by knox & holloway (1992). similar tuff layers have been observed in wells from the siri canyon, and may further be used to distinguish unit s3 from the upper part of unit s2. 0 10 20 30 40 50 60 70 80 90 100 cm 2888 m 2889 m augusta-1 2890 m sele fm bue m b fig. 36. core photographs of the lista–sele formation boundary interval in the augusta-1 well. the upper part of the bue member (lista formation) consists of mudstones with thin sandand siltstone laminae superficially resembling mudstone-in-mudstone lamination. at 2890.35 m, the bue member is overlain sharply by the laminated mudstones of the sele formation. depths are core depths. on the gamma-ray log signature. this subdivision can also be recognised on petrophysical logs from most wells in the danish sector (the subdivision is shown in five wells 49 0 10 20 30 40 50 60 70 80 90 100 cm 1689 m 1690 m nini-3 1692 m 1693 m the sele formation includes a sandstone unit (kolga member, new) in the danish north sea sector. macroand ichnofossils. fish scales and skeletal fragments are common in cores from the sele formation. bioturbation is very rare, but chondrites ispp. has been observed locally. microfossils and palynomorphs. benthic foraminifers are rare in the sele formation. the lo of an acme of the dinoflagellate genus apectodinium and the coeval lo of the shortranged a. augustum mark a level at, or a few centimetres above the base of the sele formation. the ho of a. augustum is located in the lower part of the sele formation. the ho of an influx of the dinoflagellate cerodinium wardenense marks a level in the upper part of the sele formation. the ho of an influx of the diatoms fenestrella antiqua and coscinodiscus morsianus is located in the uppermost part of the sele formation. throughout, the formation contains abundant spores and pollen, in particular pollen of the genus inaperturopollenites. depositional environment. the mudstones of the sele formation represent a mixture of pelagic fallout and dilute, low-density mud turbidites. the well-laminated character of the sediment, the high content of organic material and uranium, and the general lack of trace fossils and benthic foraminifers indicate starved sedimentation under dysoxic to anoxic bottom conditions. common diatoms indicate a high nutrient level in the water mass. the tuffs of the sele formation are evidence of extensive volcanism in the region. the significant depauperation of the benthic microfaunas during the deposition of the sele formation was most likely caused by isolation of the north sea basin (schmitz et al. 1996). the restriction and isolation of the basin was the result of a sea-level fall, possibly combined with (or caused by) tectonic uplift to the north-west (knox et al. 1981). based on microfossils, the palaeoenvironment has been suggested to represent an upper bathyal setting with a palaeodepth estimate of around 300 m (mitlehner 1996). the palynomorph assemblage indicates a marine environment characterised by a massive influx of terrestrial palynomorphs. age. sparnacian (sensu aubry et al. 2003) – early ypresian, with the lowermost level possibly of thanetian age. correlation. the sele formation corresponds to the haslund member of the ølst formation (heilmann-clausen et al. 1985). its upper part correlates with the haslund member-equivalent diatomitic knudeklint member of the fig. 37. core photographs from the nini-3 well showing the upper, arenaceous part of the sele formation. the interval belongs to the s2a subunit of the sele formation (see text for further explanation). stratigraphic position of the figured interval is shown in fig. 39. depths are core depths. fur formation in north-west jylland (danielsen & thomsen 1997). the lower boundary of the sele formation correlates with the boundary between the østerrende clay and the haslund member onshore denmark. 50 2760 2740 m 2780 s3 s2b s2a s1b s1a 2900 cleo-1 augusta-1 e-8horda fm balder fm sele fm lista fm bue mb ve mb 2880 2040 2860 2020 m m gr sonic gr sonic gr sonic cleo-1 e-8 the lower part of the sele formation, consistingof laminated, dark grey to black mudstones, correlates with the lithologically very similar, 15 m thick informal unit stolle klint clay that constitutes the lower part of the haslund member, onshore denmark (heilmann-clausen 1995). this unit is known from throughout the north sea (hardt et al. 1989; knox & holloway 1992) and constitutes most or all of the sele s1b unit of knox & holloway (1992). kolga member new member history. the kolga member consists of sandstone deposits within the sele formation. these sandstones were previously recognised by a stratigraphic working group at statoil norway in the mid-1990s and were informally referred to the hermod formation of hardt et al. (1989). derivation of name. after the goddess kolga. type well. danish sector well siri-3, 2066.4–2036.1 m mdrt (fig. 24; plate 4). reference well. danish sector well nini-3, 1717.2–1700.4 m mdrt (figs 26, 39; plate 4). distribution and thickness. the kolga member has a restricted distribution in the siri canyon in the northern part of the danish sector (fig. 20c). it reaches a thickness of up to 30 m. lithology. the member consists primarily of fine-grained to very fine-grained, olive-green to greenish grey, well-sorted, quartz-rich sandstones (fig. 39). rounded and translucent quartz grains dominate the mineralogical assemblage, but the content of glaucony grains is high (15– 25%). mica and small pyrite concretions are present in small amounts. locally, the sandstones are partly cemented by calcite and chlorite. the member usually includes one thick unit composed of amalgamated sandstone beds and a number of thinner sandstone beds interbedded with mudstones that are lithologically comparable to the sele formation mudstones described above. log characteristics. the kolga member is clearly defined on the gamma-ray log by a blocky pattern with intermediate values. this pattern differs from the high gamma-ray readings that normally characterise the lower sele formation. the kolga member may show a gradual upward decrease in gamma-ray response in its lower part (e.g. in the well nini-3; figs 26, 39). however, this does not reflect grain-size change, judging from core examination. the density log shows a blocky pattern with low density fig. 38. correlation diagram of the sele formation showing the sele units s1–3 of knox & holloway (1992). 51 v v v sele fm lista fm kolga mb bue mb grsonic 1700 1690 1710 1720 1700 1710 1720 nini-3 clay si. vf. f. m. sand c. vc. log depth core depthfig. 39. core log of the sandstonedominated kolga member encased in the sele formation mudstones in the nini-3 well. for legend, see fig. 9. intervals marked by grey bars in the core depth column are shown as core photos in fig. 37. values for the kolga member sandstones and relative high values for the interbedded mudstones (fig. 24). boundaries. the boundaries between the sandstones of the kolga member and the mudstones of the sele formation are sharp and characterised by prominent shifts on the gamma-ray and sonic logs (figs 24, 26). microfossils and palynomorphs. the kolga member is characterised by an abundance of apectodinium spp., including a. augustum. depositional environment. the sandstones of the kolga member were deposited from highly concentrated gravity flows, although the present character of the kolga member mainly reflects postdepositional liquefaction and fluidisation processes. primary sedimentary structures are common in the kolga member in some wells (e.g. sandra-1),however, indicatingthat thememberhasexperienced less postdepositional remobilisation in certain areas. age. sparnacian (sensu aubry et al. 2003) possibly including the latest thanetian. correlation. the kolga member is contemporaneous with parts of the sele formation onshore denmark. it possibly correlates with the forties sandstone member (knox & holloway 1992) in the central graben and with the teal and skadan sandstone members (knox & holloway 1992) in the southern viking graben. however, the kolga member is not contiguous with those units and has another source area. fur formation history. the fur formation is a marine diatomite with numerous ash layers. it was formally established by pedersen & surlyk (1983) with a type section in the coastal cliff knudeklint on the island of fur, denmark. its lower boundary was revised by heilmann-clausen et al. (1985). the characteristic lithology of the fur formation was subsequently recognised by thomsen & danielsen (1995; danielsen & thomsen 1997) in cuttings samples from three offshore wells located in the north-eastern part of the danish sector of the north sea, as well as in one well in the norwegian sector. 52 350 450 m k-1 gr sonic balder fm horda fm fur fm sele fm lista fm våle fm chalk group 398.9 402.2 400 m 485.5 697.1 795.3 810.7 500 600 700 800 900 dufa mb horda fm balder fm lark fm sele fm fur fm lista fm våle fm chalk gp inez-1 gr sonic type section. the coastal cliff knudeklint, the island of fur, onshore denmark (for location map, see pedersen & surlyk 1983). reference sections. silstrup south cliff, skarrehage, feggeklit, harhøj, stolleklint (for location maps, see pedersen & surlyk 1983). danish reference wells. danish sector wells k-1, 402.2– 398.9 m mdkb (fig. 40; plate 5); inez-1, 810.7–795.3 m mdkb (fig. 41; plates 1, 5). distribution and thickness. onshore denmark, the fur formation is distributed in a limited area in north-west jutland. it is c. 61 m thick in its type section. offshore, the formation occurs in a belt stretching from the northwestern coast of jutland, continuing into the norwegian sector parallel to the southern coast of norway (thomsen & danielsen 1995 text-fig. 6; fig. 42). it reaches a thickness of 15.4 m in the inez-1 well, 7.9 m in the c-1 well and 3.3 m in the k-1 well (figs 40–42). lithology. the lithology of the fur formation was described from its onshore exposures by pedersen (1981) and pedersen & surlyk (1983). it is a clayey, porous, dark grey diatomite with numerous volcanic ash layers. diatom frustules constitute 65 wt% of the rock, clay particles constitute 35 wt% (pedersen 1981). tests of coscinodiscus spp. and stephanopyxis are the major constituent of the diatomite fraction (thomsen & danielsen 1995). fine lamination is the primary sedimentary structure, but at some fig. 41. inez-1, reference well for the fur formation and type well for the dufa member. fig. 40. k-1, reference well for the fur formation. 53 25 km k-1x 3.3 15.4 7.9 inez-1 c-1 57°00' 6°00' 8°00' 56°00' 55°00' ? ? ? ? fur formation levels the lamination has been destroyed by bioturbation. the ash layers are black, graded and consist of volcanic glass particles. individual layers range from 1–20 cm in thickness but are fairly uniform in thickness over a limited area (pedersen & surlyk 1983). the diatomite recognised in cuttings samples from north sea wells by thomsen & danielsen (1995) is a similar lithology to that of the onshore sections studied by pedersen (1981) and pedersen & surlyk (1983). log characteristics. the fur formation is identified by the combination of a low gamma-ray response and low sonic readingswithinan interval of higher sonic readings characterising mudstones below and above (thomsen & danielsen 1995; danielsen& thomsen1997;figs40,41;plates 1, 5). boundaries. in the c-1 and k-1 wells, the fur formation is enveloped by the balder formation. in inez-1, the formation is bounded by the sele and balder formations (figs 40, 41; plates 1, 5). it should be noted, however, that thomsen & danielsen (1995 text-fig. 4) and danielsen & thomsen (1997 fig. 5) placed the fur formation entirely within the sele formation in the k-1 well. the discrepancy between the interpretation herein and that of the former authors is due to different interpretations of the position of the balder–sele boundary in the well. the boundary of the fur formation with the balder and sele formations is characterised by a change from laminated or structureless diatomite with ash layers to the dark mudstones of the balder and sele formations. this lithological change is reflectedonthe sonic logbyan abrupt increase in velocity (figs 40, 41; plates 1, 5). subdivision. onshore denmark, the fur formation is divided into the lower, laminated knudeklint member that contains relatively few, widely spaced ash layers and the upper, mainly structureless silstrup member with numerous ash layers (pedersen & surlyk 1983). macroand ichnofossils.themacrofossil assemblage described from onshore exposures of the fur formation encompasses fish, birds, turtles, snakes, starfish, shellfish, snails, mussels, crabs, pteropods, insects, fossil wood, leaves and fructifications (bonde 1966, 1979, 1987, 2003; pedersen 1981; pedersen & surlyk 1983; kristoffersen 2001). ichnofossils from onshore exposures include planolites ispp., teichichnus ispp., chondrites ispp. and taenidium ispp. (pedersen & surlyk 1983). microfossils and palynomorphs. diatom frustules are rockforming in the fur formation (pedersen 1981). silicofig. 42. distribution map of the fur formation with formation thickness (m) indicated for three wells. flagellates are present (perch-nielsen 1976) whereas calcareous microfossils are absent. the formation contains abundant dinoflagellates and sporomorphs (hansen 1979; heilmann-clausen 1982; willumsen 2004). depositional environment. the deposition of the fur formation diatomites took place in a long, narrow zone under upwelling conditions (bonde 1974, 1979). the upwellingwas controlled by northerly winds (bonde 1974, 1979) or it may have been created by a combination of bottom currents and bottomtopography (pedersen& surlyk 1983). age. early ypresian. correlation. onshore denmark, the knudeklint and silstrup members are largely contemporaneous with the upper part of the haslund member and the overlying værum member of the ølst formation (heilmannclausen et al. 1985), respectively, and correlate with the 54 upper part of the sele formation and the lower unit b1 (see below) of the balder formation (heilmann-clausen 1995; knox 1997 fig. 3). balder formation history. deegan & scull (1977) established the balder formation for the succession of variegated, fissile and laminated shales with interbedded tuff layers that lie between the sele and horda formations. type well. norwegian sector well 25/11-1, 1780–1705 m mdkb. danish reference wells. mona-1, 2945.0–2930.8 m mdkb (fig. 43; plate 1). siri-3, 2016.8–1998.8 m mdrt (fig. 24; plate 4). distribution and thickness. the balder formation extends over most of the central and northern north sea. in the danish sector, it reaches a thickness of more than 20 m in the siri-3 and frida-1 wells on the western part of the ringkøbing–fyn high (fig. 1) and 20 m in gwen-2 in the northern part of the danish sector of the central graben. the balder formation thins to less than 5 m towards the south-west and to less than 10 m in the eastern part of the danish sector of the north sea. the balder formation is lacking in the danish well s-1 (michelsen et al. 1998). an isochore map of the balder formation is shown in fig. 44. lithology.thebalderformationis composedof laminated, dominantly grey, fissile shales with interbedded dark and light grey, purple, buff and green sandy tuffs (fig. 45). the tuffs are normally graded and less than 5 cm thick. locally the tuff beds are slumped. the tuff layers may be cut by irregular, vertical, calcite-filled cracks up to 20 cm long (fig. 45). similar cracks have been reported from the balder formation in the grane field, norwegian sector of the north sea (haaland et al. 2000). sandstone beds, interpreted as intrusive sandstone bodies, occur locally in the balder formation. log characteristics. the balder formation is characterised by a relatively high gamma-ray values in its lower and higher parts, but shows low values in its middle part. the change in gamma-ray response is normally gradual, but relatively steep. the gamma-ray motif is mirrored by a gradual increase in sonic readings commencing at the formation base, culminating at or slightly below the level of minimum gamma-ray values in the middle part of the fig. 44. isochore map of the balder formation in the study area. the positions of the two danish reference wells, mona-1 and siri-3, are indicated in the figure. siri-3 mona-1 balder formation 5 10 15 20 thickness (m) 25 km fig. 43. mona-1, danish reference well for the balder formation. 3000 2900 m 2930.8 2945.0 horda fm balder fm sele fm lista fm våle fm vile mb ve mb bue mb chalk gp mona-1 gr sonic 55 2013 m siri-3 2012 m 2014 m 2015 m 0 10 20 30 40 50 60 70 80 90 100 cmformation, followed by a gradual decrease towards the top of the formation where the lowest sonic reading is reached. the gamma and sonic motifs together create a characteristic barrel-shaped log pattern (e.g. figs 24, 33, 38, 43; plates 1–5). boundaries. in general, the boundary with the underlying sele formation is gradational, although it can be sharp in some wells. where gradational, it is placed where the tuff layers become prominent (e.g. fig. 45). on petrophysical logs, the lower boundary is identified at a significant upward decrease in gamma-ray response accompanied by an increase in sonic readings (e.g. figs 24, 33, 38, 43). the upper boundary is at the base of the horda formation. subdivision. knox & holloway (1992) subdivided the balder formation into a lower, laminated and tuff-rich unit (b1) and a poorly laminated upper unit (b2). this subdivision can be recognised in a number of danish wells west of cecilie-1, but the b2 unit seems to be absent from danish north sea wells north-east of, and including, the cecilie-1 well. macroand ichnofossils. macrofossils have not been observed. the balder formation is non-bioturbated to moderately bioturbated. ichnofossils comprise chondrites ispp., phycosiphon ispp., planolites ispp. and thalassinoides ispp. microfossils and palynomorphs. the sele–balder boundary interval is characterised by the ho of common fenestrella antiqua and coscinodiscus morsianus (both diatoms). this event is located in the uppermost part of the sele formation but may be used as a biostratigraphic guide to locate the boundary. the diatom fenestrella antiqua characterises the balder formation and has its ho at the formation top. the dinoflagellate deflandrea oebisfeldensis shows an acme at the top of the balder formation. as observed in the underlying sele formation, the balder formation contains high numbers of spores and pollen, in particular pollen of the genus inaperturopollenites spp., and the top of the balder formation is marked by the ho of common representatives of that genus. in contrast to the overlying horda formation, calcareous benthic foraminifers are virtually absent in the balder formation. depositional environment. a restricted marine palaeoenvironment at upper bathyal depths with dysoxic to anoxic bottom conditions is suggested for the balder formation. this is based on the scarcity of calcareous microfossils and agglutinated foraminifers combined with common to abundant siliceous microfossils, especially diatoms. the fig. 45. core photographs of tuffaceous balder formation mudstones from the siri-3 well. the tuff layers are seen as light coloured, graded intervals (e.g. at 2012.42–2012.40 m). the boundary with the underlying sele formation is placed where tuffs become common, at 2015.5 m (large arrow); two tuff layers may be seen in the uppermost sele formation, at 2015.66 and 2015.90 m (small arrows). two small, lightning-shaped cracks are seen at 2015.4 m. depths are core depths. 56 2363.5 2930.8 1598.3 horda fm h2 h1 h3 l2 l3 l4 lark fm nordland gp balder fm sele fm lista fm våle fm chalk gp 1600 1700 1800 1900 2000 2100 2200 2300 2400 2500 2600 2700 2800 2900 3000 m mona-1 gr sonic fig. 46. mona-1, danish reference well for the horda and lark formations. the figure shows the tripartite subdivision of the horda formation and the l2–4 units of the lark formation. the l1 unit is absent in the mona-1 area. presence of abundant terrestrial palynomorphs further supportsa restricted,marginalmarinedepositional setting. a petrographic and geochemical study of the balder formation in the grane field, norwegian north sea sector, shows that the tuffs can be classified as representing sub-alkaline basalts and basaltic andesites of intra-plate origin (haaland et al. 2000). the tuffs are similar to the contemporaneous lower basalts in east greenland, the rockall trough and the middle series of the faeroe islands, all linked to the opening of the north atlantic (haaland et al. 2000). the volcanic phase took place at 55–52 ma. age. early ypresian. correlation. although unit b2 of knox & holloway (1992) is apparently lacking in wells in the north-eastern part of the danish sector, both units b1 and b2 can be correlated with strata onshore denmark, although unit b2 is very thin. unit b1 corresponds to the lithologically similar værum member of the widespread ølst formation onshore denmark and with the diatomaceous silstrup member of the fur formation in north-west jylland (knox 1997 fig. 3). the lower boundary of the værum and silstrup members is placed at ash layer no. +1 in the tephrachronology of bøggild (1918). the ash chronology has not been identified in the type section of the balder formation and precise correlation with the lower boundary of the værum and silstrup members is therefore uncertain. however, judging from the abundance of thick ash layers in the balder formation and the scarcity of ash layers in the underlying sele formation, it is likely that the base of the balder formation approximately correlates with ash layer no. +1, i.e. with the base of the værum and silstrup members. according to knox (1997 fig. 3), unit b2 probably correlates with the knudshoved member of the røsnæs clay formation (heilmann-clausen et al. 1985). this member has a very restricted distribution in north-west jylland where it overlies the silstrup member of the fur formation. the knudshoved member consists of a lower dark grey, pyritic clay unit rich in pyritised diatoms, and an upper greenish clay unit (heilmann-clausen et al. 1985). only a few, thin volcanic ash layers are present in the member (håkansson & sjørring 1982). based on lithological comparison, it is suggested that at least the lower, 57 pyritic part of the knudshoved member may correlate with theupper, tuff-poorunitb2 of the balder formation. stronsay group knox & holloway (1992) replaced the hordaland group of deegan & scull (1977) with two new groups: the stronsay group succeeded by the westray group (fig. 3). the two groups together comprise the light grey, green and brown coloured, soft, fissile, marine shales with thin limestone streaks that overlie the rogaland group and underlie the nordland group. these groups each contain two formations, one representing sandy shelf lithofacies and the other representing basinal mudstone lithofacies. in the central north sea, and in the danish sector, the stronsay group is represented by its mudstone facies, the horda formation (knox & holloway 1992). sandstone units of varying thickness occur at many levels in the stronsay and westray groups along the basin margin in the norwegian and british sectors, and many of these have been defined as formations or members (deegan & scull 1977; hardt et al. 1989; knox & holloway 1992). a sandstone unit also occurs in the horda formation on the ringkøbing–fyn high in the danish sector and is described here as a new member (hefring member). horda formation history. knox & holloway (1992) established the horda formation for the greenish grey basinal mudstone facies of their stronsay group that overlies the grey tuffaceous mudstones of the balder formation and underlies the greenish grey to brown mudstones of the lark formation (knox & holloway 1992). type well. british sector well 22/1-1a, 2379.5–1992 m mdkb. danish reference wells. mona-1, 2930.8–2363.5 m mdkb (fig. 46); siri-1, 2037.9–1916.5 m mdkb (fig. 47). distribution and thickness. the horda formation extends over the central and northern north sea and is present in 1916.5 2037.9 horda fm chalk gp lark fm balder fm sele fm lista fm våle fm 1900 2000 2100 2200 m siri-1 gr sonic rind mb idun mb tyr mb vile mb ve mb bue mb bue mb horda formation siri-1 mona-1 floki-1 100 200 300 400 500 600 700 800 900 thickness (m) 25 km fig. 47. siri-1, danish reference well for the horda formation. black bars show cored sections. fig. 48. isochore map of the horda formation in the study area. the positions of the two danish reference wells for the horda formation, mona-1 and siri-1, are indicated on the map. the position of floki-1, the type well for the hefring member, is also indicated. 58 gulnare-1 adda-2 alma-1 nnw sse 10 km horda fmhorda fmhorda fm l4l4l4 l2l2l2 0 500 1000 1500 2000 2500 3000 3500 twt (msec) tl tf uou tl2 th tb tc l3l3 lark fmlark fmlark fm 1000 1500 2000 2500 3000 3500 gert-1 sw ne mona-1 elna-1 sandra-1 horda fmhorda fmhorda fm h2h2 h3h3th2th2 h2 h3h3th2 10 km lark fmlark fmlark fm tl uou th tb tc twt (msec) th1th1th1 h1h1h1 fig. 49. nnw–sse-trending seismic section (rtd81-re94-17a) in the central graben showing southand eastward thinning of the horda formation. the location of the seismic section is shown in fig. 1. the l2, l3 and l4 subunits of the lark formation are indicated, as well as the mudstone-equivalent of the freja member. tl, top lark; tf, top freja; uou, upper oligocene unconformity; tl2, top l2; th, top horda; tb, top balder; tc, top chalk. fig. 50. sw–ne-trending seismic section (rtd81-rtd94-19a) showing the tripartite subdivision (h1–3) of the horda formation in the eastern part of the danish central graben (gert-1 and mona-1) and pronounced thinning of the horda formation east of the central graben (elna-1 and sandra-1). the location of the seismic section is shown in fig. 1. th1, top h1 marker; th2, top h2 marker; other abbreviations as in fig. 49. 59 all wells in the danish sector of the north sea. however, the lower part of the horda formation (fig. 4; equivalent to sequence 2 of michelsen et al. 1998) is lacking in the eastern wells r-1 and s-1 and in the eastern part of the ringkøbing–fyn high. the upper part of the horda formation (fig. 4; equivalent to sequence 3 of michelsen et al. 1998) is thin or absent in the same area (michelsen et al. 1998). the horda formation reaches a thickness of 906 m in the central graben well tordenskjold-1, but thins towards the east and south-east to less than 100 m, with minimum recorded thicknesses of 9 m in the ida-1 well and 4 m in the s-1 well. an isochore map of the horda formation is shown in fig. 48. the overall thinning of the horda formation towards the south-east, east and north-east is also shown on the seismic sections in figs 49 and 50 and on the log panel in fig. 51. lithology. the horda formation is characterised by greenish grey to greyish green fissile mudstone. subordinate limestone benches and thin layers of black mudstones occur at some levels in the formation. in many wells, particularly in the central graben, the lowermost 20–50 m of the horda formation consists of red-brown mudstones (fig. 52). this lithology is apparently lacking in the eastern wells of the danish sector. log characteristics. the horda formation is characterised by an overall stable gamma-ray and sonic log motif with a lower gamma-ray response than that displayed by the underlying balder formation and the overlying lark formation. in a few wells, the base of the horda formation shows relatively high gamma-ray values, which decrease to lower and more stable values over a short interval. the sonic readings decrease slightly upwards from the base to the top of the horda formation. boundaries. the base of the horda formation is placed at the change from the laminated, predominantly grey mudstones with interbedded sandy tuffs of the balder formation to the predominantly non-laminated, fissile, greenish grey or red-brown massive mudstones that form the basal part of the horda formation. the balder–horda boundary may be conformable or marked by a hiatus. the boundary is often difficult to pick on petrophysical logs. in basinal settings, knox & holloway (1992) advocated placing the lower boundary of the horda formation at the base of a marked gamma-ray peak believed to represent a glaucony-rich condensed layer in the basal part of the horda formation. however, in many sections in the danish sector there are two or more gamma-ray peaks in the balder–horda boundary interval. as the glauconyrich layer has not been identified with certainty in the few cores taken across the boundary in the danish sector, it is not possible to identify the key gamma-ray peak unambiguously. therefore, it is suggested that the lower boundary of the horda formation is placed on the basis of the sonic log where a gradual decrease in values in the upper part of the balder formation is succeeded by relatively stable, but somewhat lower readings in the horda formation (figs 46, 47). the upper boundary is at the base of the lark formation. subdivision. knox & holloway (1992) suggested a threefold subdivision of the horda formation (h1–3), based on lithology and biostratigraphy. a threefold subdivision can also be seen on seismic sections in the danish sector of the north sea (fig. 50). in some central graben wells, the subdivision may also be recognised on shifts in log patterns on both gamma-ray and sonic logs (fig. 51). in these wells, subtle peaks separate the three units on the gamma-ray log and coincide with the top h1 and top h2 seismic markers. based on analysis of cuttings samples, the subdivision apparently lacks lithological expression in the danish sector. knox & holloway (1992) noticed that the top of unit h1 is close to the ho of the dinoflagellate eatonicysta ursulae, and that the top of unit h2 is close to the ho of the foraminifer spiroplectammina spectabilis. this observation is supported by biostratigraphical data from the present study. a sandstone body within the horda formation has been encountered in the well floki-1 on the ringkøbing–fyn high) in the danish sector of the north sea. this sandstone is defined herein as the new hefring member (see below). macroand ichnofossils. the horda formation is moderately to intensely bioturbated. ichnofossils comprise chondrites ispp., phycosiphon ispp. and planolites ispp. microfossils and palynomorphs. in wells where the horda formation rests conformably on the balder formation, the dinoflagellate events ho deflandrea oebisfeldensis and ho dracodinium condylos occur in the lowermost part of the horda formation. characteristic microfossil datums from the lowermost part of the horda formation are the ho of the planktonic foraminifer subbotina ex gr. linaperta, which occurs abundantly, followed upwards by the almost coeval hos of the benthic foraminifers uvigerina batjesi, turrilina brevispira and gaudryina hiltermanni. a hiatus between the horda and balder formations is indicated in wells in the northern and eastern part of the danish sector by the absence of d. oebisfeldensis and d. condylos from the lower part of the horda formation. central 60 1700 1800 1900 1400 gr sonic gr sonic density gr sonic density gr sonic density 1500 1600 2000 2100 2200 2300 2400 2500 2600 2700 2800 2900 3000 3100 m 1000 1100 1200 700 800 900 1300 1400 1500 1600 1700 1800 1900 2100 2200 m 1700 1800 1900 1400 1500 1600 2000 2100 2200 2300 2400 2500 2600 2700 2800 2900 m 1800 2000 1500 2100 2200 2300 2400 2500 2600 2700 2800 3000 m kim-1 kim-1 mona-1 west lulu-3 siri-1 siri-1 1600 1700 1900 2900 2000 horda h1 h3 h2 lark l4 tl uou tl2 tl1 th th2 th1 tb l3 l2 l1 balder sele lista våle chalk grp sw ne 61 above the top of the lower, oxic part of the formation. the occasional influxes of radiolaria recorded throughout the upper part of the formation suggest that deeper marine conditions prevailed periodically. the palynofacies of the horda formation is characterised by a rich and dominant dinoflagellate assemblage with dispersed terrestrial matter (phytoclasts, spores and polfacing page: fig. 51. sw–ne-trending log panel showing eastward thinning of the horda formation. the figure also shows the variation in architecture and distribution of the horda units h1–h3 and lark units l1– 4 between the central graben (kim-1, mona-1 and west lulu-3) and the ringkøbing–fyn high (siri-1). seismic ties: tl1, top l1; th2, top h2; th1, top h1; other abbreviations as in fig. 49. fig. 52. core photographs showing red to reddish grey mudstones of the lowermost part of the horda formation in the sofie-1 well. depths are core depths. cm 1887 m 1888 m sofie-1 1901 m 1902 m 0 10 20 30 40 50 60 70 80 90 100 graben wells contain the downhole succession of the dinoflagellate cyst events ho areosphaeridium diktyoplokum, ho areosphaeridium michoudii, ho heteraulacacysta porosa and ho cerebrocysta bartonensis from the top of the horda formation, indicating an age as young as earliest rupelian (fig. 5b). in wells to the north and east, the top of the horda formation seems to be slightly older since a. michoudii is recorded from the top of the formation, indicating a mid-priabonian age (fig. 5b). significant dinoflagellate events from the middle to upper part of the horda formation are the succession of the hos of eatonicysta ursulae, diphyes ficosoides and phthanoperidinium clithridium in the middle part of the formation, and the hos of diphyes colligerum and c. bartonensis in the upper part of the formation. in central graben wells, significant microfossil events in the lower part of the horda formation include the ho of abundant radiolaria of the genus cenosphaera and the ho of the planktonic foraminifer cyclammina amplectens. key events in the middle and upper parts of the formation are the hos of pseudohastigerina spp. (planktonic foraminifers), lenticulina gutticostata, spiroplectammina amplectens and planulina costata (benthic foraminifers). the top of the horda formation contains the hos of cibicidoides truncanus and vaginulinopsis decorata. depositional environment. the lower part of the horda formation contains a microfauna that is significantly different from that of the underlying balder formation. the basal 5–40 m of the horda formation are characterised by a diverse fauna of both benthic and planktonic calcareous foraminifers together with agglutinated foraminifers. this indicates that the depositional setting was open marine, bathyal and with oxic bottom conditions. the upper part of the horda formation is characterised by an abundant and diverse agglutinated foraminifer fauna. calcareous foraminifers are very sparse or absent in this interval. the assemblage of rhabdammina discreta, haplophragmoides spp., recurvoides spp. and usbekistania charoides indicate that the upper part of the horda formation was deposited at upper bathyal depths with dysoxic bottom conditions. radiolaria occur commonly in several narrow intervals, the lowest of which is slightly 62 horda fm lark fm hefring mb våle fm balder fm sele fm lista fm chalk gp 1500 1600 1700 1800 m floki-1 1731.3 1793.4 gr sonic neutron/density fig. 53. floki-1, type well for the hefring member. len) as a minor component, indicating an open marine environment with only limited influx from surrounding terrestrial areas. age. in the central graben, where the horda formation is most complete, the formation spans from the middle ypresian (early eocene) at its base to earliest rupelian (earliest oligocene) at its top. in wells to the east and north, the top is as old as middle priabonian (see also biostratigraphic section above). this indicates that the top of the horda formation is diachronous, younging in a southwesterly direction. this is possibly due to increased erosion or longer intervals of non-deposition towards the north-east in the basin, or both. correlation. the horda formation can be correlated with the onshore danish succession of the røsnæs clay formation, the lillebælt clay formation and the søvind marl formation (heilmann-clausen et al. 1985), and the viborg formation (christensen & ulleberg 1973). the red-brown mudstones near the base of the horda formation in the central north sea can be correlated lithologically with the røsnæs clay formation and the lower part of the lillebælt clay formation. 63 the lower part of the overlying main body of greenish and greyish mudstones in the offshore succession can be correlated with the coeval and lithologically similar upper part of the lillebælt clay formation. the upper part of the horda formation can be correlated with the søvind marl formation, which consists of grey marls. the highest part of the horda formation, only observed in central graben wells, may be correlated with the viborg formation on biostratigraphic evidence. hefring member new member history. the hefring member consists of sandstone deposits within the horda formation. these sandstones have not previously been recognised as a separate unit in the danish sector. derivation of name. after the goddess hefring. type well. danish sector well floki-1, 1793.4–1731.3 m mdrt (fig. 53). distribution and thickness. the hefring member is only known from the floki-1 well located in the northern part of the danish sector. as the unit currently cannot be identified on seismic sections, its further distribution is unknown. in the floki-1 well, the member is 62 m thick. lithology. the hefring member consists of greenish grey, fine-grained, immature sandstones with glaucony grains. logcharacteristics.thehefring member ischaracterisedby aconspicuousblockysignatureonthegamma-ray, sonicand density logs (fig. 53). gamma-ray responses are lower than those of the enveloping horda formation mudstones. the hefring member can also be recognised from a combination of the density and neutron logs as the presence of pure sandstones results ina ‘cross-over’of the two logcurves (fig. 53). boundaries. the boundaries with the mudstones of the horda formation are sharp and characterised by prominent shifts on the gamma-ray and sonic logs (fig. 53). depositional environment. no cores have been taken in the hefring member, but the sandstones were probably deposited from concentrated gravity flows, based on log similarity with the other fine-grained sandstone bodies in the nearby siri canyon. age. lutetian (middle eocene) based on the age of the associated horda formation mudstones. correlation. based on biostratigraphic data, the hefring member may be contemporaneous in part with the lillebælt clay formation onshore denmark, with the lower part of the grid sandstone member (knox & holloway 1992) in the viking graben and with the upper part of the tay sandstone member (knox & holloway 1992) in the northern part of the central graben. westray group the westray group is the upper of the two groups established by knox & holloway (1992) to replace the hordaland group of deegan & scull (1977; fig. 3). in the central north sea and in the danish sector of the north sea, the westray group is represented by the lark formation. lark formation history. the lark formation was established by knox & holloway (1992) for the brownish grey mudstone-dominated lithofacies of the westray group that overlies the more variable association of red and green-grey mudstones, silty mudstones and sandstones of the horda formation and underlies the grey, sandy and shelly mudstones, siltstones and sandstones of the nordland group of deegan & scull (1977; fig. 3). the lark formation is also recognised in the danish sector although its lithology is more variable than that given in the original description. type well. british sector well 21/10-4, 1867–1217 m mdkb. danish reference wells. mona-1, 2363.5–1598.3 m mdkb (fig. 46); siri-1, 1916.5–819.3 m mdkb (fig. 54). distribution and thickness. the lark formation extends over the central and northern north sea and is probably present in the entire danish sector of the north sea. its depocentre is in the central and northern part of the danish sector, along the eastern boundary of the danish central graben, where it reaches a thickness of 1194 m in the siri-3 well. the lark formation thins west to a thickness of 389 m in the tordenskjold-1 well in the central graben, and east to a thickness of 240 m in the s-1 well on the ringkøbing–fyn high (fig. 55). 64 lithology. the lower lark formation (l1–3, see below) is dominated by dark, greenish grey, non-fissile mudstones in most wells; in some wells subordinate intervals of brownish grey mudstones are also present. thin layers of white or reddish brown carbonate are also recorded in the upper levels of the lower lark formation. the upper lark formation (l4, see below) is dominated by pale to dark brownish grey mudstones with subordinate intervals of greenish grey mudstones in its lower levels. the uppermost 50–100 m of the formation consist of yellowish grey to light brown mudstones. in eastern and northern parts of the danish sector, discrete sandstoneinterbeds and thin sandstone stringers occur throughout the formation. log characteristics. the lower part of the lark formation is characterised by an overall stable gamma-ray log signature, whereas the upper part of the formation has a more unstable signature (figs 46, 54). this change in gammaray log signature coincides approximately with the change from lithologies dominated by greenish grey mudstones to lithologies dominated by dark to light brownish grey mudstones at the base of unit l4 (see below). boundaries. the base of the lark formation is marked by a change from fissile, greenish grey mudstones of the horda formation to non-fissile, greenish grey mudstones of the lark formation. this change in lithology coincides with an abrupt increase in gamma-ray values to a consistently higher level than that displayed by the horda formation (figs 46, 51, 54). wells in the eastern part of the danish central graben and on the ringkøbing–fyn high show a conspicuous log break on the gamma-ray log at the formation boundary, whereas the log break is less pronounced in wells from the central and western parts of the danish central graben (fig. 51). although the actual increase in gamma-ray response may be limited in the latter wells, the offset is usually sharp and well defined. on the sonic log, the boundary between the horda formation and the lark formation is characterised by a transition from a stable sonic signature to one characterised by numerous fluctuations. the lark formation is overlain by the undifferentiated nordland group of deegan & scull (1977). over most of the area, the boundary seems conformable and is represented by a change from yellowish grey and light brown mudstones to medium to dark grey mudstones characte1916.5 819.3 horda fm lark fm nordland gp l4 l3 l2 l1 balder fm sele fm lista fm våle fm chalk gp 800 900 1000 1100 1200 1300 1400 1500 1600 1700 1800 1900 2000 2100 2200 m siri-1 gr sonic no log data fig. 54. siri-1, danish reference well for the lark formation. the units l1–4 are all present in this well. black bars show cored sections. 65 rised by intervals with shell-hash and coarse-grained sands. this boundary is marked by a conspicuous gamma-ray peak at the base of a 20–40 m thick interval with elevated gamma-ray values in the lowermost nordland group (figs 46, 51). this interval is further characterised by a marked double peak on the gamma-ray log. in thenorth-easternparts of the danish sector (nini-1, vanessa-1, cecilie-1 and siri-1; figs 1, 51, 54), sediments of the nordland group rest unconformably on the lark formation. in this area the uppermost lark formation and the lowermost nordland group are missing, probably due to erosion and/or non-deposition. the distinct gamma-ray peak that marks the top of the lark formation as well as the double gamma-ray peak in the lowermost nordland group are lacking in these wells, and therefore the top of the lark formation is more difficult to identify on petrophysical logs. subdivision. the lark formation can be subdivided into four major mudstone packages, l1–l4, based on seismic and log evidence (figs 46, 49, 51, 54, 56; plates 1–5). these units are described below; isochore maps of the units are shown in fig. 57a–d. l1 (figs 51, 56a, b, 57a; plates 1, 4, 5) this unit has been recognised in the north-eastern part of the danish sector only (fig. 57a). it is bounded beneath by the th marker and above by the tl1 marker (fig. 56a, b). it is characterised by downlapping reflectors and represents a south-westwards prograding mudstone succession. on the gamma-ray log, the l1 unit is characterised by a relatively high and relatively stable response. in most wells, it shows a weakly concave pattern, going from a relatively high gamma-ray response at its base, over a gamma-ray low halfway through the unit to a level close to starting level at the top of the unit (e.g. ida-1, inez-1, k-1, f-1 and sandra-1; plates 1, 4, 5). in the siri-1 and siri-3 wells, near the south-western limit of the l1 unit, the gamma-ray log motif instead appears slightly convex (figs 51, 54; plates 1, 4). the l1 unit consists predominantly of greenish grey mudstones but also includes yellowish brown and dark grey mudstones. l2 (figs 46, 51, 56a, b, 57b; plates 1–5) the unit is recognised over the entire study area. on the gamma-ray and sonic logs the unit is characterised by a stable log signature. the gamma-ray log shows two to three slightly concave patterns with signatures similar to that of the l1 unit (fig. 51; plates 1, 4). the lithology is characterised by dark beige-grey to greenish grey mudstones, greenish colours becoming dominant towards the top of the unit. l3 (figs 46, 51, 56a, b, 57c; plates 1–5) this unit is encountered in the northern and eastern parts of the danish sector, east of the central graben (fig. 57c) but is not recognised on logs or seismic sections in the central graben area. it is characterised by stable gamma-ray and sonic log signatures (figs 46, 51; plates 1–4). the unit consists almost invariably of dark, greenish grey mudstones. l4 (figs 46, 51, 56a, b, 57d; plates 1–5) the unit is recognised over the entire study area (fig. 57d). the interval is characterised by a slightly more unstable gamf-1 inez-1 siri-1 francisca-1 frida-1 tordenskjold-1 s-1 mona-1 25 km 200 300 400 500 600 700 800 900 1000 1100 1200 1300 freja member distribution dufa member distribution lark formation thickness (m) fig. 55. isochore map of the lark formation in the study area. the positions of the two danish reference wells, mona-1 and siri-1, are indicated in the figure. the map also shows the distribution of the sandstones of the dufa and freja members and the location of their type and reference wells, inez-1 and f-1, and francisca-1 and frida-1, respectively. 66 fig. 56. a: sw–ne-trending seismic section (rtd81-re94-22a) showing the complex architecture of the lark formation and its subdivision into l1–4 units. the vertical white line indicates change in section direction. the locations of the two seismic sections are shown on fig. 1; abbreviations as in figs 49 and 51. 0 500 1000 1500 2000 2500 3000 3500 0 500 1000 1500 2000 2500 3000 3500 tordenskjold-1 f-1 tabita-1 cecilie-1 siri-3 sw ne horda fmhorda fm l2l2 lark fmlark fm l4 horda fm l2 aa l2l2 l1 lark fmlark fm l4 l3 l1 10 km lark fm lark fm l4 l3l3l3 l4 l3 uou tl2 th tb tc tl tf uou tl2 tl uou tl2 tl1 th tb tc twt (msec) twt (msec) 67 25 50 75 100 125 thickness (m) l1 25 km 25 km 100 200 300 400 thickness (m) l2 100 200 300 400 500 600 700 800 thickness (m) 100 200 300 400 500 600 700 800 thickness (m) l3 l4 a b c d fig. 57. isochore maps of lark formation subunits. a: l1. b: l2. c: l3. d: l4. a and d are at the same scale, b and c are at the same scale. 1500 2000 2500 3000 3500 twt (msec) 20 km nini-2 d-1 nolde-1 nw se tl uou tl2 tl1 tb tc l1 l2l2 l3 l4 l1 l2 l3 l4 bb fig. 56. b: nw–se-trending seismic section (rtd81-re94-14a) showing subdivision of the lark formation and marked thinning of this formation towards the south-east. the locations of the two seismic sections are shown on fig. 1; abbreviations as in figs 49 and 51. 68 ma-ray and sonic log signature than that of the underlying units (figs 46, 51; plates 1–5). it is dominated by brown to yellowish brown mudstones, but in some wells an interval of greenish grey mudstones occurs in its lower part. in wells to the east and north, thin sandstones are interbedded with the mudstones and become more frequent towards the top of the unit. two thick sandstone units occur in the lark formation on the ringkøbing–fyn high and are described here as two new members (dufa and freja members). macroand ichnofossils. only observed in cores taken in the freja member (see below). microfossils and palynomorphs. farthest to the north and east the basal part of the formation includes the downhole succession of ho areosphaeridium diktyoplokum and ho a. michoudii indicating a late priabonian (late eocene) age for the base of the formation in this area. in the central graben area, the base of the lark formation is significantly younger. here it contains an event succession characteristic of the middle and lower rupelian (lower oligocene) stage (hos of phthanoperidinium amoenum, achilleodinium biformoides and phthanoperidinium comatum). the top of the lark formation is bracketed by a number of conspicuous biostratigraphic events: the uppermost part contains the hos of the benthic foraminifers asterigerina staeschei and elphidium inflatum followed downhole by the ho of uvigerina tenuipustulata. dinoflagellate events near the top of the lark formation include the hos of apteodinium spiridoides and cousteaudinium aubryae. the lowermost part of the overlying nordland group contains the hos of the calcareous microfossils bolboforma clodiusi, bolboforma spiralis and bolboforma metzmacheri, the ho of the benthic foraminifer bulimina elongate, and the ho of the dinoflagellate cyst cannosphaeropsis passio. a large number of hos characterise the lark formation; key events are listed in fig. 5c. depositional environment. the l1 unit is characterised by abundant agglutinated foraminifers dominated by rhabdammina discreta and similar tubular taxa, together with haplophragmoides spp. and recurvoides spp. the microfaunal assemblage indicates that the unit was predominantly deposited in an open marine, dysoxic palaeoenvironment at upper bathyal depths. the l2 unit and most of the l3 unit are characterised by an increasing abundance and diversity of calcareous benthic and planktonic foraminifers. the relative proportions of agglutinated, calcareous planktonic and benthic foraminifers vary considerably from well to well, indicating pronounced lateral changes in the depositional environment. the calcareous plankton/benthos ratio is usually low, indicating a neritic setting for most of the succession, but in a few restricted intervals it may reach 1:2 or even 1:1 indicating an outer neritic setting. thus, the foraminifer assemblage indicates an open marine, neritic to outer neritic setting with well-oxygenated bottom conditions for the lower to middle part of the lark formation. the microfossil assemblage in the uppermost part of the l3 unit as well as the l4 unit is dominated by calcareous benthic foraminifers, and agglutinated foraminifers are generally rare. epifaunal and shallow infaunal foraminifers are more common than deep infaunal taxa, indicating oxic bottom conditions during this interval. in general, the microfaunal assemblage in this part of the lark formation suggests that it was deposited in a neritic, probably middle neritic, palaeoenvironment over most of the study area. the palynofacies assemblage in the lark formation is characterised by a rich dinoflagellate assemblage and abundant dispersed terrestrial matter (phytoclasts, spores and pollen), indicating an open marine environment with considerable influx from nearby land areas. stratigraphic variations in the relative abundance of terrestrial palynomorphs in the lark formation suggest successive pulses of progradation and backstepping of the palaeocoastline. age. the lark formation is of priabonian to serravallian (late eocene to middle miocene) age with eocene sediments being present in the l1 unit only. the l1 unit is priabonian to early rupelian (early oligocene) in age, the base of the unit being oldest farthest to the north and east and younging towards the south and west. the age of the l2 unit is rupelian; the l3 unit is rupelian in age in its lower part and chattian (late oligocene) in its upper part. the rupelian–chattian boundary is located in the lower part of the unit. the chattian–aquitanian (oligocene– miocene) boundary is located just above the top of the l3 unit. in some wells, a hiatus is indicated at this level by the clustering of hos. the chattian–aquitanian, aquitanian–burdigalian, burdigalian–langhian and langhian–serravallian stage boundaries are all located in the l4 unit. the uppermost part of the lark formation is of midserravallian age. correlation. based on biostratigraphic correlation, the lowermost l1 unit is probably largely coeval with the viborg formation onshore denmark, and with sequence 4.1 of michelsen et al. (1998). the l2 unit may be corre69 lated with the linde clay onshore denmark (informal mudstone unit described by heilmann-clausen 1995). intervals in the l3 unit may be correlated with the branden clay (ravn 1906) onshore denmark, based on lithological similarities and biostratigraphy. intervals in the lark formation around the l3–l4 boundary (around the chattian–aquitanian boundary) may be correlated with the two lowermost, clay-rich units of the onshore vejle fjord formation(thebrejningclayandvejlefjordclayoflarsen & dinesen1959).theuppermostpart of the lark formation possibly correlates with the onshore arnum formation (sorgenfrei 1958) and the hodde formation onshore denmark (rasmussen 1961), based on biostratigraphy. dufa member new member history. the dufa member comprises a thick sandstonedominated unit that occurs within unit l3 of the lark formation in the northern and eastern part of the danish sector of the north sea. the unit has not been previously described. derivation of name. after the goddess dufa. type well. danish sector well inez-1, 697.1–485.5 m mdkb (figs 41, 58; plates 1, 5). reference well. danish sector well f-1, 337.5–324.3 m mdkb (figs 58, 59; plate 5). distribution and thickness. the dufa member is present in the north-eastern part of the danish sector of the north sea (fig. 55). in its type well, the dufa member is 210 m thick and consists of three major sandstone units with thicknesses 30–120 m (fig. 41). the sandstone units are separated by mudstone intervals up to 20 m thick. towards the north, in the f-1 well, the lower sandstone units are fig. 58. n–s-trending composite seismic section (rtd81-re94-45/rtd81-re94-09). blue-coloured lines indicate the outline of the dufa member. the gamma-ray logs from the inez-1 and f-1 wells are inserted (see figs 55 and 59 for depth-converted gamma-ray logs for the two wells). the horda formation is thin in this area and the top horda reflector (th) is therefore indistinguishable from the top balder reflector (tb). the location of the seismic section is shown in fig. 1; abbreviations as in fig. 49. inez-1f-1 0 500 1000 twt (msec) tf tl2 tb/th n s dufa mb dufa mbdufa mb 5 km 70 well, the lower sandstone intervals of the dufa member show the presence of a number of 5–10 m thick sandstone packets showing blocky, decreasing-upwards gamma-ray log signatures suggesting coarsening-upwards sand bodies. these sandstones are separated by intervals of fig. 60. francisca-1, type well for the freja member. black bars show cored sections. 1400 1500 1600 1700 1800 m 1562.8 1840.7 freja mb lark fm francisca-1 gr sonic missing (figs 58, 59). the dufa member is absent in wells west of f-1. lithology. the lower sandstone units predominantly consist of coarsening-upwards successions of very fine-grained to fine-grained, greenish brown, muddy sandstones. the upper sandstone unit fines upwards and consists of mediumto coarse-grained, quartzitic, relatively pure sand with intervals rich in glaucony. lignite has been observed in cuttings samples. log characteristics. the member is characterised by an overall blocky signature on the gamma-ray log. in the type fig. 59. f-1, reference well for the dufa member. dufa mb lark fm 300 400 500 600 m f-1 324.3 337.5 gr sonic 71 mudstones with higher gamma-ray response. the upper unit is characterised by an overall blocky signature with minor gamma-ray peaks and trends suggesting a number of fining-upwards intervals, 10–20 m thick, and a few coarsening-upwards intervals, 5–10 m thick (fig. 41). boundaries. the upper and lower boundaries of the member with the mudstones of the lark formation are sharp and characterised by prominent shifts on the gamma-ray log (figs 41, 59). depositional environment. judging from seismic evidence, the dufa member is positioned partly on the offlap break, partly seaward of it (fig. 58). based on this palaeosetting and the presence of lignite in cuttings samples, the dufa member sandstones are interpreted to represent deltaic, shallow-marine sediments, probably deposited in pulses during an overall relative sea-level low. age. rupelian, based on the age of the enveloping mudstones. correlation. there are no danish onshore correlatives to the dufa member. the correlation with the norwegian offshore successions is currently uncertain. freja member new member history. the freja member is a conspicuous sandstonedominated unit that occurs within the upper levels (l4) of the lark formation in the northern and central parts of the danish sector of the north sea. the unit has not previously been described. derivation of name. after the goddess freja. type well. danish sector well francisca-1, 1840.7–1562.8 m mdrt (figs 60, 61). reference well. danish sector well frida-1, 1623.5–1487.7 m mdrt (fig. 62; plate 4). distribution and thickness. the freja member is present in the northern and central parts of the danish sector of the north sea (fig. 55). in its type well, the freja member spans a stratigraphic interval of c. 280 m and includes major sandstone units separated by subordinate intervals of mudstones (fig. 60). in the cecilie-1 well, the member is c. 150 m thick whereas in the frida-1 well the member attains c. 130 m (fig. 62). lithology. in its type well, the lower half of the freja member consists of very fine-grained to fine-grained quarzitic sandstones with many thin mudstone interbeds. the member becomes less muddy in the upper third of this interval. the upper third of the freja member consists largely of relatively pure quarzitic, very fine-grained sandstones with mudstone interbeds becoming frequent towards the top (fig. 60). between these two major sandstone units is a c. 60 m thick interval dominated by mudstones but with carbonate-cemented, sandstone-dominated packets in its upper part. log characteristics. the freja member has an overall blocky gamma-ray log signature. in the type well, its lowermost part (1840.7–1750 m) can be split into a number of smaller units with blocky or increasing-upwards gamma-ray log signatures separated by gamma-ray peaks. in comparison, the overlying sandstones (1750–1720 m) display a more stable, low gamma-ray log pattern with few gamma-ray log spikes (fig. 60). the mudstone-dominated interval (1720–1660 m) separating the two sandstone-dominated units in the type well generally shows high gamma-ray values: thin calcite-cemented sandstone packets are intercalated with the mudstones in this interval (e.g. 1680– 1670m)and show decreasing-upwards gamma-ray values. boundaries. the lower boundary of the freja member with the lark formation mudstones is sharp and characterised by prominent shifts on the gamma-ray and sonic logs. in the type well, where the upper levels of the freja member are characterised by interbedded mudstones and sandstones, the upper boundary of the member is less prominent. in this well, it is placed at the top of the uppermost discrete sandstone bed, at 1562.8 m (fig. 60). macroand ichnofossils. intervals with shell debris have been observed in core sections of the freja member in the francisca-1 well. ichnofossil genera from the freja member comprise chondrites ispp., phycosiphon ispp., planolites ispp., terebellina ispp., thalassinoides ispp. and zoophycos ispp. depositional environment. the freja member represents stacked successions of thickand thin-bedded turbidite sands deposited in submarine channels and proximal levee environments (figs 60, 62). the upper parts of the turbidite successions show transitions from normally graded turbidites, deposited in slightly more distal levee environ72 francisca-1 w e n s 2 km 0 500 1000 1500 2000 2500 3000 twt (msec) freja memberfreja memberfreja member horda fmhorda fm lark fmlark fm lark fmlark fm horda fm lark fm l3l3 l2l2 l4l4 lark fm tl tf uou tl2 th tb ments and minor turbidite channels, to mainly silty turbidite deposits that represent distal levee and fan fringe environments and the transition to the open slope. the source of the sand was probably a marginal marine shelf environment, judging by the abundance of the marginal marine acritarch paralecaniella indentata. age. in the type well, the freja member is chattian to aquitanian in age, based on the age of mudstones within and bounding the member. in the frida-1 well, the freja member is entirely chattian in age. correlation. the freja member is broadly contemporaneous with the vejle fjord formation onshore denmark, with the vade formation (hardt et al. 1989) in the norwegian central graben and with the skade formation (hardt et al. 1989) in the viking graben. thick, coarsening-upward sandstone bodies are present above the dufa memfig. 61. composite seismic section (dk1–5623a re94/dk1–0448b re94) with the freja member indicated between the reflectors uou and tf. the gamma-ray log from the francisca-1 well is indicated on the figure (see fig. 60 for depth-converted gamma-ray log). the vertical white bar indicates where the section changes direction. the location of the seismic section is shown on fig. 1; abbreviations as in fig. 49. 73 1487.7 1623.5 freja mb lark fm 1400 1500 1600 1700 m frida-1 gr sonic ber in the inez-1 well (shown as unnamed sandstones in fig. 2); these sandstones may be contemporaneous or even contiguous with those of the freja member. however, confident correlation on the basis of log and seismic data is not possible at present. acknowledgements aage bach sørensen (geus) is thanked for help with seismic interpretations. yvonne desezar, johnny e. hansen and birthe amdrup are thanked for preparation of microfossil and palynology samples. the referees robert w.o’b. knox (british geological survey) and paul van veen (conocophilips norway) are thanked for their constructive criticism of the manuscript; the editorial contributions of adam a. garde, jon r. ineson and martin sønderholm are gratefully acknowledged. this work was made possible through grants from the danish energy authority, under the energy research project framework 2000. fig. 62. frida-1, reference well for the freja member. black bar shows cored section. references ahmadi, z.m., sawyers, m., kenyon-roberts, s., stanworth, c.w., kugler, k.a., kristensen, j. & fugelli, e.m.g. 2003: paleocene. in: evans, d. et al. 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(eds): petroleum geology of the continental shelf of north-west europe: proceedings of the 2nd conference, 3–39. london: heyden & son ltd. 78 79 80 preface nielsen 2023: geus bulletin 56. 8356. https://doi.org/10.34194/geusb.v56.8356 1 of 2 preface troels f.d. nielsen1* 1department of mapping and mineral resources, geological survey of denmark and greenland (geus), copenhagen, denmark *correspondence: tfn@geus.dk received: 19 jun 2023 accepted: 19 jun 2023 published: 22 dec 2023 keywords: cryptic variation, cumulates, layered intrusion, gabbros, mineral chemistry, phase layering abbreviations: dlc: danish lithosphere centre geus: geological survey of denmark and greenland geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: jakob kløve keiding (geus, denmark) reviewed by: none. preface to: thy et al. 2023. (this volume) funding: none. competing interests: none. additional files: none. the palaeogene skaergaard intrusion, east greenland, has since its first description by wager & deer (1939) been a foremost natural laboratory for the study of low-pressure fractionation of basaltic melt. ocean floors are composed of basalt and the processes that control compositions of basaltic melts are fundamental to the dynamics of the earth. this special issue of geus bulletin by peter thy, christian tegner and charles e. lesher is the most recent in a more than eighty year succession of trendsetting works on the evolution of the skaergaard intrusion and evolution of basaltic melt. the early sample collections, housed in the universities of oxford and cambridge, were used for the development of fractionation models, mapping of the distribution as well as the partition of major and trace elements between melts and liquidus phases in basaltic melts. these and many other studies resulted in the monumental “layered igneous rocks”, edited by l.r. wager and g.m. brown (wager & brown 1968). by the 1970s, the original collection in the oxford university museum of natural history and the sedgwick museum of earth sciences, university of cambridge had seen extensive use. the science developed and new and more detailed sampling was now required. fieldwork and sampling in the later part of the 20th century led to a flurry of new studies by more research groups, including neil irvine of carnegie’s geophysical laboratory and the university of oregon group lead by alex mcbirney. despite all the efforts, no consensus was reached on the many fundamental chemical and physical processes in basaltic magma chambers. the available collections did not provide sufficiently systematic and detailed information for the modelling of the evolution of the basaltic melt and the genesis of the precious metal deposit that had been discovered in the intrusion. exploration drill cores offered the possibility for tight stratigraphic sampling through >1000 m of gabbro. in combination with surface sampling, a new standard profile with superior systematics could be established. in the summer of 2000, the danish lithosphere centre (dlc) organised fieldwork and the transport of the drill cores to the museum of natural history in copenhagen. this special issue is part of the legacy of the dlc and a fulfilment of the then envisaged potential of the skaergaard intrusion for fundamental studies of processes in basaltic magma chambers. thy et al. (2023, this volume) analyse and compile in unprecedented detail the variation in major and trace element compositions of bulk-rock samples and electron microprobe compositions of rock-forming minerals in the layered series on the floor of the intrusion. in addition, they include a wealth of information on all samples from the skaergaard intrusion that are housed in museum, university and survey collections in denmark, the uk https://doi.org/10.34194/geusb.v56.8356 https://orcid.org/0000-0002-4932-3869 mailto:tfn@geus.dk geusbulletin.org nielsen 2023: geus bulletin 56. 8356. https://doi.org/10.34194/geusb.v56.8356 2 of 2 and us, topographic and geological maps, aerial photographs in archives and images of thin sections used in the study. with all its data and appendices, the monograph is the most comprehensive collection of petrological information for the skaergaard intrusion and the layered series on the floor of the intrusion, ever to have been made available for research. the detailed modelling presented re-evaluates petrogenetic constraints and tests petrogenetic models in the literature. the modelling is based on liquidus proportions established in experimental studies of appropriate melt compositions. thy and co-authors conclude that evolution observed in the gabbros of the layered series is predominantly controlled by crystal fractionation of bulk liquid. late in the evolution and above the boundary between upper zone a and b, the gabbros crystallised from ponded fe-rich, immiscible melt. redistribution of granophyric melts and decoupling of included and excluded trace elements, not only in bulk liquid but also in mush melts, complicates the modelling of trace element distributions in the late and evolved stages of crystallisation in the crystal mushes of the intrusion. this special issue is a treasure trove for all who study the intrusion and are searching for a data set to be matched with experimental and numeric modelling of low-pressure fractionation of basaltic melt. with this contribution, the skaergaard intrusion continues to be a foremost natural laboratory for modelling of melts formed in solidifying basaltic magma chambers. troels f.d. nielsen emeritus, geus, denmark references thy, p., tegner, c., & lesher, c.e. 2023: petrology of the skaergaard layered series. geus bulletin 56, 8327 (this volume). https://doi. org/10.34194/geusb.v56.8327 wager, l.r. & deer, w.a. 1939: geological investigations in east greenland: part iii. the petrology of the skaergaard intrusion, kangerdlugssuaq, east greenland. meddelelser om grønland 105, 352 pp. wager, l.r. & brown, g.m. 1968: layered igneous rocks, 588 pp. edinburgh & london: oliver & boyd. http://www.geusbulletin.org/ https://doi.org/10.34194/geusb.v56.8356 https://doi.org/10.34194/geusb.v56.8327� https://doi.org/10.34194/geusb.v56.8327� preface references sørensen et al. 2025: geus bulletin 59. 8380. https://doi.org/10.34194/3cy30452 1 of 29 research article the kangâmiut dyke swarm in west greenland: a new map and insights into their tectonic evolution kai sørensen*1 , john a. korstgård2 , flemming mengel3 , willy lehmann weng1 1geological survey of denmark and greenland (geus), copenhagen, denmark; 2department of geoscience, aarhus university, aarhus, denmark; 31795 view point rd, boulder, co, usa abstract this contribution presents a new map of the palaeoproterozoic kangâmiut dyke swarm in central west greenland. the map is based on publicly available aerial imagery, the scale and quality of which allowed for quick and efficient interpretation across a large area. the kangâmiut dyke swarm has played a pivotal role in the identification and characterisation of the southern margin of the nagssugtoqidian orogen. change in dyke orientation from nne-trending in the south to ene-trending farther north is accompanied by increasing deformation in both dykes and host rocks. the zone where dykes and host rocks are totally parallelised defines the southern structural and metamorphic front of the nagssugtoqidian orogen. we document variable changes in orientation of the dykes and their density to estimate the crustal extension accompanying dyke emplacement. the average width of the 123 dykes is 25 m (80% are <50 m). these dykes occur with an average frequency of 3.4 dykes per km and make up 6–11% of the outcrops. these data reveal subordinate groups of dykes with ese and ne orientations and track regional changes. at present, their ages relative to the dominant nne-trending swarm are not known. the swarm generally extends from south of maniitsoq northwards to the ikertooq shear zone; however, we identified features north of the ikertooq shear zone, which we speculatively interpret to represent the northernmost occurrence of the kangâmiut dyke swarm. the tectonic consequences of this interpretation – if correct – allow us to estimate the amount of shortening across the ikertooq shear zone during the nagssugtoqidian orogeny to be more than 150 km. if the other major tectonic boundaries in the orogen, including the nordre strømfjord shear zone, were the loci of similar shortening, the current extent of the orogen may represent only a fraction of pre-nagssugtoqidian crust in central west greenland. *correspondence: kaisoerensen@hotmail.com received: 08 aug 2024 revised: 10 dec 2024 accepted: 12 dec 2024 published: 04 jul 2025 keywords: kangâmiut dykes, palaeoproterozoic, nagssugtoqidian orogen, crustal shortening, west greenland abbreviations: cno: central nagssugtoqidian orogen dlc: danish lithosphere centre ggu: grønlands geologiske undersøgelse isz: ikertooq shear zone nno: northern nagssugtoqidian orogen sno: southern nagssugtoqidian orogen snf: southern nagssugtoqidian front td: total depth geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: thomas find kokfelt (geus, denmark) reviewed by: jeroen van gool (independent researcher, denmark); kyle mayborn (western illinois university, usa) funding: see page 27 competing interests: see page 27 additional files: see page 27 1. introduction in the preface to their 1961 geological survey of greenland bulletin, arne noe-nygaard and hans ramberg described the pioneering work in the precambrian basement of west greenland. they wrote: “in 1946 the geodetic survey was able to deliver us blue-prints of the newly made 1:250 000 scale map of the holsteinsborg area, and consequently we set to work here, this region being as suitable as anything else on the west coast.” (noe-nygaard & ramberg 1961, p. 4). that report contained the first extensive map of the precambrian rocks between qeqertarssup tunua in the north (69°n) and the nuuk region in the south (63°45n). the region chosen turned out to be so much more than ‘suitable’ because it contains a swarm of basic dykes – the kangâmiut dykes – which become increasingly deformed and metamorphosed when traced from the south towards the north and ne through kangerlussuaq. this remarkable feature had already been described by hans ramberg as early as 1948 based on fieldwork in 1946–1947. he wrote: “when one is sailing in through the long fiord kangerdlugssuaq, it is amazing to see how the originally ‘young’ cross-cutting kangâmiut diabases (fig. 2) gradually begin to take part in the plastical deformation of the gneisses” (ramberg 1949, p. 316). he further described https://doi.org/10.34194/3cy30452 https://orcid.org/0009-0009-8938-092x https://orcid.org/0000-0003-2143-8514 https://orcid.org/0000-0001-5028-9998 https://orcid.org/0000-0002-9083-7642 mailto:kaisoerensen@hotmail.com https://creativecommons.org/licenses/by/4.0/deed.ast sørensen et al. 2025: geus bulletin 59. 8380. https://doi.org/10.34194/3cy30452 2 of 29 geusbulletin.org how the diabases are folded, recrystallised into amphibolites, affected by metasomatism and broken into large boudins. the term “nagssugtoqidian” also made its first appearance in ramberg (1949) in his description of the nassuttooq fjord as long, branched and situated in the middle part of the folded chain of post-kangâmiut diabase age. in the fjord, he described “the most excellent exhibitions of folds and other tectonic deformations” and designated this precambrian folded chain the “nagssugtoqides” (ramberg 1949, p. 318). ramberg (1949) divided the “nagssugtoqides” into three main complexes: the “egedesminde”, the “isortoq” and the “ikertoq” gneiss complexes (fig. 1), according to their degree of regional metamorphism. the central isortoq gneiss complex is located in granulite facies, and the bordering ikertoq and egedesminde gneiss complexes are located in amphibolite facies of presumed nagssugtoqidian age. it is clear that the kangâmiut dykes played a crucial role in the early identification of what we now call the nagssugtoqidian orogen and in the unravelling of the nature of the border zone between the ‘older’ southern block and the ‘younger’ northern complexes. in this contribution, we briefly review the previous research in the nagssugtoqidian in west greenland, and present the results of a recent photo interpretation effort of the kangâmiut dyke swarm. we discuss their orientation and thickness variations, and finally, present a model for the present geographic distribution of the dykes. 1.1. early work on the nagssugtoqidian orogen in 1965, the geological survey of greenland (grønlands geologiske undersøgelse, ggu) started systematic reconnaissance mapping in the area around and north of kangerlussuaq with the aim of completing the earlier coastal mapping reported in noe-nygaard & ramberg (1961). escher (1966) found a major tectonic discordance between the amphibolite facies rocks of the ikertoq gneiss complex and those of the granulite facies of the isortoq gneiss complex (fig. 1). this zone of faulting and shearing was interpreted as a nw-dipping thrust zone, now termed the ikertooq shear zone (isz). this reconnaissance work led to the publication of the 1:500  000 geological map covering the kangerlussuaq–nuussuaq region (“søndre strømfjord–nugssuaq region”, escher 1971) and the recognition of the alternating pattern of ‘linear or straight belts’ and areas without a strong preferred orientation (escher 1970) which is a key characteristic of the nagssugtoqidian of west greenland. reconnaissance in the area south of 66°45’n (see fig. 2) was carried out in the late 1970s and included a major part of the kangâmiut dykes (see allaart & jensen 1979) and led to the subsequent publication of the 1:500 000 geological map covering sioqqap sermia – kangerlussuaq (“frederikshaab isblink – søndre strømfjord”; allaart 1982). along with more recent maps, including the compilation by garde & marker (2010), these two maps are now part of the seamless 1:500 000 geological map of greenland (pedersen et al. 2013) and depict the extent of the kangâmiut dyke swarm including the generally held view of a northern limit of the swarm along the isz. in addition to the work by ggu, two projects by aarhus university (denmark) and the university of liverpool (uk) and a multi-institution project by the danish lithosphere centre (dlc) led to several detailed studies of nagssugtoqidian geology. a major result of the work by the aarhus group, initiated in 1965 by bondesen (1966), was the recognition of a major shear zone along the nassuttooq fjord (bak et al. 1975a). the completed 1:100 000 “agto” (attu) map sheet 67 v.1 nord (olesen 1984; location in fig. 3) was another major result (olesen et al. 1979). as a result of the dlc work, the 1:100 000 ussuit map sheet 67 v.2 nord was published (van gool & marker 2004; location in fig. 3). a tectonic model of the nagssugtoqidian orogen resulting from the work by dlc is summarised in van gool et al. (2002) and van gool & marker (2007). the liverpool precambrian boundary programme (watterson 1974) aimed to determine the displacements fig. 1 original overview map of the nagssugtoqidian from ramberg (1949), with their place name spellings, showing the proposed subdivision of the precambrian rocks between kangaamiut (kangâmiut in the figure) and qasigiannguit (christianshaab in the figure) in west greenland. the isortoq gneiss complex is in granulite facies, while the adjoining ikertoq and egedesminde gneiss complexes are in amphibolite facies. slightly modified from ramberg (1949). 50 km n https://doi.org/10.34194/3cy30452 http://www.geusbulletin.org/ sørensen et al. 2025: geus bulletin 59. 8380. https://doi.org/10.34194/3cy30452 3 of 29 geusbulletin.org 50 km archaean palaeoproterozoic undi�erentiated arfersior�k quartz diorite sisimiut charnockite archaean gneisses variably re-worked during the nagssugtoqidian orogeny supracrustal rocks ma�c dykes kangâmiut ma�c dyke swarm (schematic) 52° 67° 68°68° 67° 52°54° 50° 48° 69° 66° 54° sisimiut kangaamiut sermiat torsuup nunaa s n f greenland ice sheet attu qeqertarsuup tunua qasigiannguit ilulissat aasiaat kangaamiut so ut he rn a rc ha ea n fo re la nd kangâmiut-1 maniitsoq nordre strøm�ord shear zone nordre isortoq shear zone ikertooq shear zone a s s c n o n n o sn o a iso rto q kangerlussuatsiaq ikertooq nassu tto oq sisim iut isortuat qorlortoq kangerlu ssu aq itillip ilua eqalussuit tasiat kangerluarsussuaq simiutaq nunakerlua inuppaat quuat itilleq n fig. 2 simplified geological map of west greenland, showing the internal subdivisions of the nagssugtoqidian orogen (sno: southern nagssugtoqidian orogen, cno: central nagssugtoqidian orogen, nno: northern nagssugtoqidian orogen, after marker et al. 1995). the kangâmiut dykes investigated in this study occur from 65°10´n to 67°45´n. north of itillip ilua and the southern nagssugtoqidian front (snf), country rocks are largely archaean orthogneisses and minor supracrustal rocks, which are all variably to completely reworked during nagssugtoqidian orogenesis. the thin stippled line at kangaamiut sermiat traces the aujassoq shear zone (“aujassoq-evighedsfjord shear belt” of allaart & jensen 1978). ‘kangâmiut-1’ is the location of the well drilled in 1976 by the french oil company total. black triangles: thrust faults. inset map shows the extent of the nagssugtoqidian orogen across southern greenland. slightly modified from korstgård et al. (2024) and includes information from escher (1971) and allaart (1982). https://doi.org/10.34194/3cy30452 http://www.geusbulletin.org/ sørensen et al. 2025: geus bulletin 59. 8380. https://doi.org/10.34194/3cy30452 4 of 29 geusbulletin.org and history of the boundary region (watterson 1974, p. 33) and focused on the western part of the isz (fig. 2). a joint paper by the aarhus attu mapping group and the liverpool group discusses tectonic implications of the shear zones observed in the nagssugtoqidian (bak et al. 1975b). in an important paper, kalsbeek et  al. (1987) suggested, based on isotopic and elemental geochemistry, that the arfersiorfik quartz diorite in the eastern part of the nassuttooq region (see fig. 2) originated in a subduction-related oceanic arc setting, thus introducing fig. 3 index map showing the locations of figures, photos and place names cited in the text. extents of aerial photos and maps are shown with the prefix ‘fig’. location of field photographs and other displays are shown by the figure number. red lines: trends and lengths of transects a–f are shown schematically. grey stippled lines: simplified major tectonic boundaries. snf isz nisz nssz fig. 20a fig. 16 fig. 17 fig. 25a fig. 13 fig. 15 fig. 12a fig. 22 25b fig. 11a fig. 4 fig. 21 fig. 11b fig. 5 map sheet agto 67 v1n map sheet ussuit 67 v2s 30 24 26a 29 26b 23 10a6b 6a 10b 10c f a d e b c maniitsoq kangaamiut kangerlussu atsia q tasersiaq isorto q arfersior�k sisimiut itilleq attu inussuttusup tunua kangaamiut sermiat torsuup nunaa kangerl ussu aq sisimiut isortuat nassuttooq ikertooq itillip ilua maligiaq sarfartoq sarfannguit nunaat tunu saqqap kangerluarsua geodetic reference: wgs84 >1500 m 300–1500 m 0–300 m basemap: geus, g2.5m geus, ww, 2021-09-08 projection: utm zone 84 68°n 67°n 66°n 68°n 67°n 66°n 54°w 52°w 52°w 50°w 50°w n 0 25 km https://doi.org/10.34194/3cy30452 http://www.geusbulletin.org/ sørensen et al. 2025: geus bulletin 59. 8380. https://doi.org/10.34194/3cy30452 5 of 29 geusbulletin.org plate tectonic concepts in earnest into the interpretation of nagssugtoqidian geology. a more detailed tectonic model involving the arfersiorfik quartz diorite and accompanying lithologies in the same area was proposed by van gool et al. (2002). more recently glassley et al. (2007, 2014) presented evidence of ultra-high pressures in rocks near nassuttooq. 1.2. previous research on the kangâmiut dykes the results of the first phase of regional mapping were compiled in 1956 and published in two map sheets, scale 1:500  000 (noe-nygaard & ramberg 1961). in these maps, observations of dykes and other rocks and structural features were generally restricted to coastal areas. figure 4 shows a detailed view of the original map covering the southern part of the area reported by noenygaard & ramberg (1961) and illustrates the general density and orientation of the dykes. the continued reconnaissance mapping by ggu in the 1960s and 1970s of the area between sioqqap sermia (62°30’n) and kangerlussuaq (66°45’n) led to the publication of a 1:500 000 map (allaart 1982), which was the first depiction of the kangâmiut dykes as a regionally continuous swarm with a consistent geometry. the reconnaissance mapping combined with interpretation of aerial photos also led to the publication of the seminal paper on the reorientation of the kangâmiut dyke swarm by escher et  al. (1975). their generalised map is reproduced here (fig. 5). a sharp transition from undeformed to deformed and re-orientated dykes is described to occur along a continuous boundary between the kangaamiut sermiat ice cap and the greenland ice sheet and this was termed the ‘main nagssugtoqidian boundary’ (here labelled snf in fig. 5). in contrast, it was noted that a much more gradual transition is seen w and nw of kangaamiut sermiat. 10 km n fig. 4 original reconnaissance map from noe-nygaard & ramberg (1961), with their place name spellings. location in fig. 3. diagonal shading with diamond symbols: agmatitic hypersthene-gneiss and pyroxene-gneiss of enderbitic to quartz-dioritic composition. black lines: observed kangâmiut dykes. purple shading: ultrabasic rocks (lower right of map). yellow shading: granulite facies metasediments (lower right of map). slightly modified from noe-nygaard & ramberg (1961). https://doi.org/10.34194/3cy30452 http://www.geusbulletin.org/ sørensen et al. 2025: geus bulletin 59. 8380. https://doi.org/10.34194/3cy30452 6 of 29 geusbulletin.org it was observed during the early reconnaissance that in addition to the dominant nne-trending dykes that appear throughout the region, there appears another set of dykes, also deformed and metamorphosed during the nagssugtoqidian (escher 1970). these dykes trend between e and ese and were referred to as ‘ese dykes’. for reasons described next, we use the term kangâmiut dykes for both sets. scattered dykes of trends deviating from the nne and ese directions occur throughout the region but are not numerous. as part of the liverpool project, detailed mapping of dykes was carried out by jack (1978) in an area near inussuttusup tunua (fig. 3) and by nash (1979) in an area around the mouth of itillip ilua (fig. 3). also, as part of the liverpool project, korstgård (1979, 1980) mapped dykes based on aerial photos and field work in the ikertooq to itilleq region (fig. 3), thus extending the escher et al. (1975) map (fig. 5) northwards to include the dyke geometry in the isz and its vicinity. korstgård (1979, 1980) also described the metamorphic and structural changes of the dykes along a transect from itillip ilua to maligiaq in inner ikertooq fjord (fig. 3). along this transect, a transition from low amphibolite facies in the south to granulite facies in the north was documented. adjacent to the ‘main nagssugtoqidian boundary’ of escher et al. (1975, p. 160), the authors described a transition zone, 10–20 km wide characterised by the presence of “many shear belts affecting both the east– west and kangâmiut swarms.” in these belts, escher et  al. described brown granulite facies gneisses retrogressed to pale biotite gneisses, and dolerites converted to amphibolites. it is worth emphasising here, that the granulite facies gneisses are of archaean age, whereas the retrogression to amphibolite facies and the metamorphism of the dykes are the result of palaeoproterozoic (nagssugtoqidian) events. the main nagssugtoqidian boundary ne of kangaamiut sermiat (fig. 5) was investigated during the dlc project. fieldwork in 1994 seemed to indicate that the southern part was “… formed by a few nw– nnw-dipping ductile thrusts arranged in a sinistral en echelon array…” hageskov (1995b, p. 14). following 0 10 20 30 40 50 km tasersiaq sarfartoq itilleq s n f itillip ilua kangerlussu aq tasersuaq kangaamiut sermiat tasersiaq fig. 4 sarfartoq kangerlussuatsia q umiiviit inussuttusup tunua nunakerlua torsuup nunaa saqqap kangerluarsua n fig. 5 original map of kangâmiut dykes from escher et al. (1975). location in fig. 3. the frame in the sw corner shows the area covered by fig. 4. snf: southern nagssugtoqidian front. the trace of the snf from escher et al. (1975) and mapping by hageskov (1995a, 1995b) and olsen (2000). modified from korstgård et al. (2024). https://doi.org/10.34194/3cy30452 http://www.geusbulletin.org/ sørensen et al. 2025: geus bulletin 59. 8380. https://doi.org/10.34194/3cy30452 7 of 29 geusbulletin.org fieldwork in 1995, hageskov (1995a, p.10) concluded that the southern nagssugtoqidian boundary zone includes at least three “en echelon thrusts which from north to south are named the taserssuaq, the paradise and the sarfartoq thrust.” hageskov also suggested that the individual thrusts cannot have accommodated significant displacement, and observed that some granitic rocks near the greenland ice sheet occurred within the footwall as well as within the hanging wall. as a part of the dlc project, olsen (2000) mapped the vicinity of the sarfartoq to further refine the delineation of the nagssugtoqidian ‘structural front’ in that area and showed how multiple thrust segments could be identified. thrusts mapped by hageskov (1995a, 1995b, 1997, 1998) and olsen (2000) are shown in fig. 5. engström & klint (2014) mapped areas near the head of kangerlussuaq and while their emphasis was on lineaments of potential interest for the hydraulic properties of the gneisses, they did describe the ‘nagssugtoqidian structural front’ to be a complex array of thrust faults. across the nagssugtoqidian boundary from south to north, the nne dykes show a clockwise rotation, whereas the ese dykes display an anti-clockwise reorientation. this observation was used by escher et al. (1975) as a tool to identify an overall nnw–sse-orientated direction of shortening during deformation (escher et al. 1975, fig. 7). the intrusive relations of the two sets of dykes in the inussuttusup tunua area were described by escher et al. (1975, 1976) and jack (1980) who speculated that the two swarms were emplaced syn-kinematically along conjugate shear fractures with the nne and ese dykes being largely contemporaneous (watterson 1974; escher et al. 1976). in contrast, korstgård et al. (2024) and the present contribution envisage an overall extensional regime for the emplacement of the kangâmiut dykes. in the region between maniitsoq and itillip ilua (fig. 2), mengel et  al. (1996) distinguished three sets of dykes based on mutual intersections and inferred that in this area, the e–w dykes are oldest, nne-trending dykes are younger, and the main swarm of nne–ne-trending dykes is the youngest. however, the only available age determinations are from the youngest dykes (see section 1.3), so how much the emplacement of these sets are separated in time remains unknown. the highest density of kangâmiut dykes is found across a c. 140 km section from the head of kangerlussuatsiaq to the mouth of itillip ilua (fig. 5) where, according to escher et al. (1975), more than 100 dykes wider than 15 m occur and represent a crustal widening of 2–3% (escher et al. 1975, p.158). within the central part of the nne swarm, allaart & jensen (1979) found a dyke density reaching up to 7% between kangaamiut and the mouth of kangerlussuaq. to address these significantly varying densities, we have also assessed dyke densities in a few key locations within the main swarm (section 3.2). the western limit of the kangâmiut dyke swarm is clearly not exposed, but a petroleum exploration well drilled c. 100 km west of the mouth of kangerlussuaq by the french oil company total in 1976 (66°10n, 56°14w, fig. 2) may well provide some constraints. the kangâmiut-1 well was drilled to a subsea depth of 3874 m in 180 m deep water. among the material collected, were what appeared to be remnants of a kangâmiut dyke, recovered near total depth (td). sørensen (1977) described the material collected, and in the report noted that “with the exception of subcalcic clinopyroxene all of the primary features of the kangâmiut dykes on-land nne ese nne fig. 11a a b 2 km 500 m n n fig. 6 aerial photographs of kangâmiut dykes. locations in fig. 3. a: aerial photo f1946. kangâmiut dykes hosted by amphibolite facies gneisses in kangerlussuaq c. 57 km from the mouth of the fjord. width of the fjord sw of the delta lobe is 2.4 km. the frame indicates the extent of the outcrop shown in fig. 11a. b: aerial photo e3116. kangâmiut dykes in granulite facies host gneiss showing a clear intersection between nne and a later ese dyke, north shore of kangerlussuaq, 20 km from the mouth of the fjord. the ese-trending dyke is 80 m wide. https://doi.org/10.34194/3cy30452 http://www.geusbulletin.org/ sørensen et al. 2025: geus bulletin 59. 8380. https://doi.org/10.34194/3cy30452 8 of 29 geusbulletin.org 68°n 67°n 66°n 68°n 67°n 66°n 54°w 52°w 52°w 50°w kangerl ussu aq sisimiut itilleq kangaamiut maniitsoq attu kangerlussuaq kangaamiut sermiat sisimiut isortuat nassuttooq isortoq kangerlussu atsia q kangerluarsussuaq tasersiaq inussuttusup tunua simiutaq torsuup nunaa itillip ilua saqqap kangerluarsua ikertooq inuppaat quuat nssz nisz itz geodetic reference: wgs84 >1500 m 300–1500 m 0–300 m basemap: geus, g2.5m geus, ww, 2021-09-08 projection: utm zone 84 0 25 km n kangâmiut dykes (schematic) structural trends in transition zone structural trends in shear/thrust zones fig. 7 major nagssugtoqidian structures and generalised trends of dykes based on interpretation of aerial photos. this map is a simplified version of supplementary file s1. red lines: structural trends in the major shear and thrust zones. black lines: general trend of dykes where observable as continuous, intrusive bodies. green lines: gneiss trends in the area south of the ikertooq shear zone (isz), where gneiss structures and highly deformed dykes are parallelised. nssz: nordre strømfjord shear zone. nisz: nordre isortoq shear zone. https://doi.org/10.34194/3cy30452 http://www.geusbulletin.org/ sørensen et al. 2025: geus bulletin 59. 8380. https://doi.org/10.34194/3cy30452 9 of 29 geusbulletin.org are met with on this site, but on-land many dykes are likewise without primary pyroxene. the full range of secondary features, except for the most extreme, secondary deformation is represented within the material, but the overall picture is of an undeformed kangâmiut dyke, both between levels 3874 and 3849 and the fragments in the side wall cores. the metamorphic grade of the dykes indicates low amphibolite facies, with bluish green amphibole and secondary plagioclase of oligoclase composition coexisting with epidote” (sørensen 1977, p. 34). the potential significance of this observation is discussed further in section 4. despite the common assumption that kangâmiut dykes do not occur north of the isz, there are many published accounts of mafic dykes in this area with preserved discordances that bear a strong resemblance to kangâmiut dykes. a few of these occurrences are summarised next. several examples were documented across the attu (agto) map sheet covering the western half of nassuttooq (fig. 3), including those occurring in the vicinity of attu (figs 2, 3) and south of eqalussuit tasiat described by sørensen (1970, 1971) and glassley & sørensen (1980) and those centrally in the map sheet north of nassuttooq (mengel 1983). during the mapping in the area immediately n and ne of attu, árting (2004) encountered several mafic metadykes forming a minor ese-trending swarm. farther east, metadolerites have also been identified in a gneiss panel in the nordre strømfjord shear zone exposed in inner arfersiorfik fjord (sørensen et al. 2006, fig. 3a). these are affected by nagssugtoqidian metamorphism but locally preserve intrusive relationships. in gneisses in the qorlortoq valley (fig. 2), van gool & marker (2007) describe the presence of numerous dykes; however, later work (mcintyre et al. 2021) notes that despite the field similarities, trace element compositions suggest that the qorlortoq dykes are likely unrelated to the kangâmiut dykes and are presumably archaean in age. detailed descriptions of field occurrence, geochemistry, petrology and petrogenesis of the kangâmiut dykes are presented in several papers, among them windley (1970), jack (1978), korstgård (1979, 1980), zeck & kalsbeek (1981), bridgwater et al. (1995), mengel et al. (1996), mayborn (2000), cadman et al. (2001), mayborn & lesher (2004, 2006), mayborn et al. (2008) and most recently by korstgård et al. (2024). 1.3. regional geological framework the nagssugtoqidian orogen can be traced from west to east across greenland (e.g. bridgwater & gormsen 1968; bridgwater 1976; escher & watt 1976; bridgwater et al. 1989; kalsbeek 1989; van gool et al. 2002; review in nutman et  al. 2008; kolb et  al. 2016) and is part of a collage of palaeoproterozoic orogens forming the northern margin of the archaean north atlantic craton (e.g. bridgwater et  al. 1973). farther west, in northern labrador, the palaeoproterozoic torngat orogen occurs along the western margin of the archaean nain craton (e.g. korstgård et al. 1987; hoffman 1989). east of greenland, the palaeoproterozoic laxfordian orogen in nw scotland borders the archaean lewisian complex (e.g. park 2022). the similarities between these coextensive palaeoproterozoic orogens in labrador, greenland and scotland have been pointed out by many (e.g. bridgwater et al. 1989; park 1995; mason & brewer 2004; st-onge et al. 2009). based on the work outlined in sections 1.1 and 1.2, the following framework describes the nagssugtoqidian orogen in west greenland in terms of structural, metamorphic and geochronological evolution of its component parts. marker et  al. (1995) subdivided the orogen into the southern, central and northern nagssugtoqidian orogen (sno, cno and nno, respectively, fig. 2). these are broadly similar to ramberg’s (1949) ‘ikertoq’, ‘isortoq’ and ‘egedesminde’ gneiss complexes (fig. 1), but are separated by major tectonic breaks of shear or thrust zones. each subdivision is composed of archaean rocks variably deformed and metamorphosed during nagssugtoqidian orogenesis and each also contains layers of supracrustal rocks of both archaean and palaeoproterozoic ages (marker et  al. 1999). as mentioned in section 1.2, kangâmiut dykes are generally reported to be restricted to the sno and the southern archaean foreland. the southern limit of the sno is the ‘main nagssugtoqidian boundary’ described in section 1.2, but the term ‘southern nagssugtoqidian front’ (snf) is preferred here. this marks the zone north of which the kangâmiut dykes and their host gneisses are deformed and rotated into the penetrative ene structural grain characteristic of the nagssugtoqidian (figs 2, 5). the northern boundary of the sno is the northerly-dipping isz, in which kangâmiut dykes and their host gneisses are deformed into parallelism. this zone separates granulite facies in the north from amphibolite facies rocks in the south and contains metasedimentary panels, strongly deformed along with the gneisses. the cno extends northwards up to the nordre strømfjord shear zone (bak et al. 1975a), a major transcurrent structure extending from the coast to the greenland ice sheet. the cno includes the central nordre isortoq shear zone, which is characterised by several large metasedimentary packages. the cno contains two juvenile calc-alkaline intrusive bodies, the arfersiorfik quartz diorite in the east and the sisimiut charnockite in the west (fig. 2). both have been dated at c. 1.9 ga and are interpreted to represent arc magmatism related to https://doi.org/10.34194/3cy30452 http://www.geusbulletin.org/ sørensen et al. 2025: geus bulletin 59. 8380. https://doi.org/10.34194/3cy30452 10 of 29 geusbulletin.org subduction during continent convergence prior to collision (e.g. kalsbeek et al. 1987; glassley et al. 2010). to the north, the nno is not as strongly reworked as the sno and cno. its northern boundary is poorly defined as a gradual transition to the rinkian fold belt farther north (e.g. escher & pulvertaft 1976; grocott & pulvertaft 1990; van gool & marker 2007; grocott & mccaffrey 2017; grocott et al. 2023). most of the archaean gneisses give published ages around 2.8 ga and there is geochronological evidence of a major crust-building episode at 2870–2810 ma and subsequent reworking at 2790–1680 ma (kalsbeek & nutman 1996; connelly & mengel 2000; connelly et al. 2006). zircons from kangâmiut dykes intruding archaean rocks (near kangaamiut and itillip ilua, fig. 2) have been dated to 2045–2021 ma (kalsbeek & nutman 1996; nutman et al. 1999; connelly et al. 2000; nilsson et al. 2019). the high-grade processes associated with nagssugtoqidian orogeny, specifically collision, thrusting and transcurrent shearing, occurred c. 1870–1780 ma. deformation and thrusting across the snf occurred during the earlier part of this interval, while shearing and thrusting in the main tectonic zones in the orogen took place 50–100 myr later (e.g. korstgård et al. 2024). it is notable, that related arc magmatism (kalsbeek et  al. 1987) preceded the early phases of collision by up to c. 50 million years (kalsbeek & nutman 1996; connelly et al. 2000; van gool et al. 2002). the extent of nagssugtoqidian metamorphic overprints in the southern foreland has been discussed by mengel et al. (1995, 1996), mayborn & lesher (2006) and korstgård et al. (2024). 2. new mapping 2.1. data the primary source of data for the new map of the kangâmiut dykes is a collection of aerial photos from the p0886 survey available from the greenland portal (https://maps.greenmin.gl/; geus 2023). the aerial survey was acquired in 1985 from high altitude (14 000 m) and presented at 1:150  000 scale. the availability of these images has made it possible to investigate large areas within a more reasonable time frame than would have been possible with the pre-existing aerial photos in 1:40 000 scale. the map produced as part of the liverpool project (korstgård 1979, fig. 1) was based on the interpretation of the 1:40 000 data set and used c. 200 photos to map an area of c. 4000 km2, while the new map presented here based on the 1:150  000 data set involved interpretation of c. 125 photos to cover an area of c. 30 000 km2. the scale cited for the aerial surveys applies to the scale of the contact copies made from the photographic plates. the 1:40  000 scale aerial photos were also used by escher et al. (1975) in their mapping (fig. 5). in addition to these aerial photos, 1:20 000 eox sentinel-2 cloudless satellite images (eox 2018) and google earth satellite images (https://earth.google.com) were used. these images are in colour, which can be helpful in the identification of dykes. while the available eox scenes are of good quality and cover the entire area, they are of restricted resolution. the google earth scenes are of variable quality, but in some areas, they are both of good quality and better resolution than the eox data. mapping was done by unaided visual inspection, without using a stereoscope. the transfer of observations from the aerial photos to the map was greatly facilitated by the availability of high quality, shaded topographic maps with many lakes and rivers shown (available on the greenland portal). all aerial photos shown here are from the p0886 survey and are referred to by their serial numbers from the greenland portal, without the prefix ‘p0886’. 2.2. mapping objectives the photo interpretation strategy follows that employed in korstgård (1979, 1980), in which two linear map features were targeted, both of which were confirmed by field observations where possible. one feature is lineaments that reflect map traces of gneiss and the second feature is mafic dykes. these are especially recognisable where they occur as individual discordant dykes in less deformed areas but can be harder to distinguish in strongly deformed zones where they are parallel with the dominant gneissic structures. dykes hosted by the darker granulite facies rocks are typically less distinct than those in lighter amphibolite facies gneisses (e.g. figs 6a, b). in addition, in many inland areas, the degree of exposure is generally poor and here, dykes are more difficult to distinguish. in the main dyke swarm, where major dykes occur at a typical frequency of c. 3 per km, only a selection of dykes has been mapped, while outside the main swarm ‘coverage’ is higher. combined with the highly varying nature of exposure and quality of images, this means that the map shows the structural trends of the dykes, not their varying density or thickness. the results of this mapping are shown in full at scale 1:500 000 in supplementary file s1 and a simplified version, scaled down for publication here, is shown in fig. 7. the extent of the dykes has been subdivided into 12 areas (fig. 8), in which areas 1–5 represent the general south to north changes in the dyke swarm, while the remaining areas illustrate the slightly more complex variations east and south of areas 1–5. the orientation data collected during this work are summarised in rose diagrams https://doi.org/10.34194/3cy30452 http://www.geusbulletin.org/ https://maps.greenmin.gl/ https://earth.google.com sørensen et al. 2025: geus bulletin 59. 8380. https://doi.org/10.34194/3cy30452 11 of 29 geusbulletin.org 68°n 67°n 66°n 68°n 67°n 66°n 54°w 52°w 52°w 50°w 50°w area 5a area 5b area 6 area 7 area 8 area 9 area 10a area 10b area 10c area 11 area 12 area 1 area 2 area 3 area 4 itillip ilua kangaamiut maniitsoq arfersiorfik sisimiut attu inussuttusup tunua kangerlu ssu aq sisimiut isortuat ikertooq nssz nisz itz snf nassu tto oq fig. 25a fig. 25b isortoq itilleq tasersiaq simiutaq sarfannguit nunaat kangerlussuatsia q geodetic reference: wgs84 0 25 km >1500 m 300–1500 m 0–300 m basemap: geus, g2.5m geus, ww, 2021-09-08 projection: utm zone 84 n fig. 8 map showing the extent of areas 1–12 with rose diagrams of dyke orientations for each area. grey stippled lines: simplified major tectonic boundaries. nssz: nordre strømfjord shear zone, nisz: nordre isortoq shear zone, isz: ikertooq shear zone, snf: southern nagssugtoqidian front. https://doi.org/10.34194/3cy30452 http://www.geusbulletin.org/ sørensen et al. 2025: geus bulletin 59. 8380. https://doi.org/10.34194/3cy30452 12 of 29 geusbulletin.org for each area (fig. 8). azimuths of dykes are grouped in 10-degree bins for each area and lengths were summed and represented as percentages of cumulative length. 3. description of the dyke map the area studied covers the central and southern part of the nagssugtoqidian orogen along with its southern foreland (figs 2, 7, 8). from south to north, the kangâmiut dykes are progressively deformed and re-orientated to variable degrees: in the southern archaean foreland the dykes occur as steeply inclined, nne-striking, little to moderately deformed continuous bodies. the mapping reported here speculates that dykes with similar geometry occur north of the isz in the area shown in fig. 7 (see section 3.1.1). within the isz (fig. 2), strongly deformed dykes occur in 100–1000 m wide panels of gneisses interleaved with metasediments (10–100 m) in which no dykes are found (see section 3.3). throughout the sno, the magnitude of deformation is highly variable. in this study, we call this the ‘transition zone’. note that this is quite different from the transition zone defined by escher et al. (1975, fig. 2). 3.1. dykes in the southern foreland and area 5 3.1.1. areas 1–4 in areas 1–4, from south towards north (fig. 8), and ne along kangerlussuaq, a gradual change in orientation of the kangâmiut dykes can be observed. from the southernmost coastal part of area 1 to the coastal part of area 3 just south of itilleq (figs 8, 9), the kangâmiut dykes change strike from c. 30° to c. 55° (fig. 9). the transition from nne to ene strike can also be seen from the mouth of kangerlussuaq and farther ne along the fjord. while most dykes in areas 1, 2 and 3 are subvertical, ne along kangerlussuaq, the change in strike is accompanied by a change in dip (fig. 10). this was first described by ramberg (1949), and likewise we interpret this gradual change in both strike and dip as a result of progressive deformation. 3.1.2. age relations between nne and ese dykes it has been suggested (e.g. by escher et  al. 1975) that the ese dykes predate the main swarm of nne– ne-trending kangâmiut dykes. however, escher et al. (1976, p. 533) observed that intersections showed no consistent age relations in the inussuttusup tunua area, suggesting that the two dyke sets are effectively of the same age. we have made similar observations, including an example of an ese dyke cutting an nne dyke (fig. 6b), and localities where both relationships are seen (fig. 11a, b), supporting the claim made for the inussuttusup tunua area (escher et al. 1976). we interpret these observations to indicate that nne and ese dykes intruded broadly simultaneously, and use the term ‘kangâmiut dykes’ to describe both sets of dykes. age data reported in section 1.3 came from the main nne–ne-trending swarm, and at the time of writing none of the ese dykes have been dated. several studies (e.g. hamling et al. 2010) have found that the emplacement of dykes can modify the stress field (especially the minimal compressive stress), such that later dykes preferentially intrude with different orientations. we think something similar may account for the varying orientation of the kangâmiut dykes. 3.1.3. area 5 kangâmiut dykes have not been mapped – or previously reported – in areas 5a and 5b, north of the isz. this area is entirely underlain by granulite facies gneisses and supracrustal rocks with an overall ene-trending structural grain (figs 2, 7). in this study, we have identified what we interpret to be a population of mafic dykes in both areas 5a and 5b (figs 12 and 13). the orientation of these dykes is predominantly nne in fig. 9 rose diagrams for areas 1, 2, 3 and 4 demonstrating gradual clockwise change in orientation of the nne dykes towards the ‘typical’ nagssugtoqidian ene trend. location of areas 1–4 shown in fig. 8. within each area, azimuths of dykes are grouped in 10-degree bins and lengths of dykes are summed and represented as percentages of cumulative length. 20 40 60 80 100 120 140 160 102030% 20 40 60 80 100 120 140 160 102030% 20 40 60 80 100 120 140 160 102030% 20 40 60 80 100 120 140 160 102030% area 1 area 2 area 3 area 4 n ew s n ew s n ew s n ew s https://doi.org/10.34194/3cy30452 http://www.geusbulletin.org/ sørensen et al. 2025: geus bulletin 59. 8380. https://doi.org/10.34194/3cy30452 13 of 29 geusbulletin.org both areas with a small number of ese dykes (fig. 14), similar to the distribution in the foreland south of the isz. we speculate that these may be contemporaneous with the kangâmiut dykes south of the isz. it should be emphasised at the onset of this discussion, that the presence and nature of these dykes have not been confirmed on the ground and are only interpreted from aerial imagery. a corollary of this interpretation is that this area between the nagssugtoqidian nordre isortoq shear zone and isz (fig. 2) locally escaped the penetrative deformation that characterises much of the cno, which would require a radical re-interpretation of the nature and extent of nagssugtoqidian deformation. it is noteworthy, that the westernmost interpreted dykes do not overlap with the eastern extent of the sisimiut charnockite as would be expected and as shown on the 1:500 000 geological map (garde & marker 2010). the nne dykes in the eastern part of area 5b (figs 12 and 13) display a density comparable to that of the nne-trending kangâmiut dykes in the coastal area between kangaamiut and itilleq (figs 2, 7). the densest central part of the inuppaat swarm occurs in the area south and sw of the inuppaat quuat glacier (fig. 12b) and extends as far as c. 30 km west of the ice margin (figs 7, 12a), possibly coinciding with the prominent terrain lineaments visible on the eox satellite data (shown on fig. 12a). dykes have also been interpreted immediately north of the isz (see example in. fig. 15). 3.1.4. areas 11, 12 and the aujassoq shear zone in this section, we consider the structural significance of the topographic lineaments that follows maniitsup sermilia (indicated by ‘1’ in fig. 16), ‘elephant foot valley’ (‘2’ in fig. 16, informal name) and lake tasersiaq (‘3’ in fig. 16). a shear zone (described variably as a fault, mylonite zone or shear zone) forming a prominent topographic lineament from maniitsoq through area 1 and continuing into elephant foot valley (fig. 16) was observed during ggu reconnaissance mapping in 1976–1978 and was termed the aujassoq fault zone (allaart et al. 1977) or aujassoq-evighedsfjord shear belt (allaart & jensen 1979). it corresponds to the evighedsfjord shear zone of bak et al. (1975b). in this zone, granulite facies gneisses are retrogressed to amphibolite facies. most of the zone of presumed most intense deformation is not exposed but it is visible locally along the shores of maniitsup sermillia and kangerlussuatsiaq (figs 16 and 17) and has been documented by notes on several field maps resulting from ggu’s reconnaissance mapping from the 1970s. the most complete trace of the aujassoq shear zone, as we refer to the structure here, is depicted on the field compilation map in scale 1:100 000 by allaart et  al. (1977). the retrogression and associated fabric development in maniitsup sermilia (fig. 16) was noted during reconnaissance by b. windley (1964–1967). gneisses and dykes exposed along elephant foot valley display a strong, valley-parallel structural grain (figs 17a, b). on the escher et al. (1975) map, the dykes are shown as occurring across the entire width of elephant foot valley (fig. 5). our aerial photo observations suggest that the valley-parallel shear zone is up to 3 km wide. fig. 10 field photographs showing the three stages – a to b to c – in the progressive deformation of the kangâmiut dykes as seen in outcrops along the south shore of kangerlussuaq. locations of photos in fig. 3. a: the arrow points to the peak named ‘stoppenålen’, the nw face of which (left side in image) is formed by a nne dyke dipping c. 80° nnw. viewing direction is ene. b: dykes strike ne and dip c. 50 degrees to nw. viewing direction ne. c: dykes strike ene and dip c 45 degrees to the nnw. viewing direction ene. in all three photos, the peaks are 1000–1300 m above sea level. photos courtesy of simon mose thaarup, taken in 2016. a b c https://doi.org/10.34194/3cy30452 http://www.geusbulletin.org/ sørensen et al. 2025: geus bulletin 59. 8380. https://doi.org/10.34194/3cy30452 14 of 29 geusbulletin.org the aujassoq shear zone may continue into area 12 (figs 8, 16) as two or three strands striking ene and e along and north of lake tasersiaq (fig. 16). the entire area 12 is characterised by variably orientated dykes and our observations are mostly of rather short dyke segments due to the nature of exposure. the rose diagram (fig. 18) therefore conveys the dyke geometry better than the map (fig. 7). a major part of the identified dykes is the collection of ese dykes that have been deformed into parallelism with the above-mentioned strands of the aujassoq shear zone (fig. 16). a few nne dykes were also identified, but these were mainly outside the zones of deformation. escher et al. 1975 (p. 159) noted that post-dyke shear belts occur in a few isolated places within the southern zone near the eastern part of lake tasersiaq (boellstorff 1968, referenced in escher et al. 1975). they remarked that most of these shear zones are parallel with and adjacent to the wider dykes of both swarms and result in amphibolitization of dyke margins and retrogression fig. 12 dykes in the eastern part of area 5b (location in fig. 8). these dykes are informally named the inuppaat swarm after the inuppaat quuat glacier visible in the figure. a: overview of the area. stippled white line is a major lineament referred to in the text. eox image. box shows the extent of fig. 12b. b: aerial photo l0840 showing an enlarged view of several large dykes of the inuppaat swarm. dykes that are readily visible at this scale are indicated by arrows. widest dyke is 120 m. inuppaat quuat inuppaat quuata b 1 km5 km n n fig. 12b dyke 1 dyke 3 250 m500 m n na b dyke 1 dyke 3dyke 2 1 2 fig. 11 examples of mutually intersecting dykes. a: ese-dyke cutting (at ‘1’) and being cut (at ‘2’) by nne dykes. cliff face is c. 1.3 km wide. width of the widest dyke c. 50 m. location in figs 3, 6a. source: aerial photo f1946. b: ese dyke (dyke 2) cutting as well as being cut by nne dykes (dyke 1 and dyke 3). image is 1 km wide. dykes are 50–60 m wide. location in fig. 3. source: google earth. fig. 13 n–nne-trending dykes south of the meltwater plain isortup kuua. location in fig. 3. e–w width of image is 23 km. the nature of exposure does not permit an assessment of dyke density. only in a few places can a clear colour contrast provide confirmation of a dyke, but terrain lineaments nevertheless suggest the occurrence of a substantial number of dykes. source: aerial photo h927. 5 km n isortup kuua https://doi.org/10.34194/3cy30452 http://www.geusbulletin.org/ sørensen et al. 2025: geus bulletin 59. 8380. https://doi.org/10.34194/3cy30452 15 of 29 geusbulletin.org of the country rock. in this area, garnets have been observed in the kangâmiut dykes (e.g. loewe et al. 1962) indicating the extent of nagssugtoqidian metamorphic effects in the foreland region south of the snf. despite the aujassoq shear zone being a major structural feature, its relative age is currently not known with certainty. our observations suggest an overall parallelism between dyke orientations and the structural grain in the host gneisses, whereas allaart & jensen (1979) describe the dykes as truncating the gneiss structures in this area. a possible scenario is that the shear zone is an original archaean feature, and that this anisotropy partially controlled the orientation of the kangâmiut dykes. the area was then thermally overprinted during the nagssugtoqidian. 3.1.5. area 10a in area 10a, north of tasersiaq, both nwand ne-trending dykes were identified (fig. 19). these give rise to a set of distinct, dyke-parallel topographic lineaments. the relationship between these two sets of dykes is not known and no observations of dyke intersections were made. 3.1.6. areas 10b and 10c in areas 10b and 10c – part of the southern archaean foreland – the ese dykes are most abundant (figs 7, 20) and form a minor swarm, which seems broadly parallel to the ese dykes found farther west (north and nw of tasersiaq, fig. 5). this eastern part of the ese swarm dominates close to torsuup nunaa (location in fig. 3), and we will term it the torsuup swarm. the dense part of the swarm is c. 15 km wide. across this width we observed some 50 major dykes (along the n–s extent of fig. 20a), thus indicating a dyke density comparable to that characterising the central part of the major nne swarm farther west (e.g. fig. 5). 3.2. dyke densities in the southern archaean foreland and in area 5 the overall highest density of nne dykes is found across a c. 100 km transect from the head of kangerlussuatsiaq to the mouth of itillip ilua fjord (location in area 5b area 5a n ew s 20 40 60 80 100 120 140 160 102030% n ew s 20 40 60 80 100 120 140 160 102030% fig. 14 rose diagrams showing dyke orientations in areas 5a and 5b. locations in fig. 8. rose diagram construction described in fig. 9. n 1 km fig. 15 nne–ne-trending dykes immediately north of the eastern extension of the ikertooq shear zone (isz). limited colour contrast between dykes and gneisses suggests that host rocks are in granulite facies. a few dykes are indicated with arrows. the southern part of the image shows the e–w trend of the isz. location in fig. 3. source: aerial photo k0349. https://doi.org/10.34194/3cy30452 http://www.geusbulletin.org/ sørensen et al. 2025: geus bulletin 59. 8380. https://doi.org/10.34194/3cy30452 16 of 29 geusbulletin.org fig. 3), but as noted earlier, the detailed density – and its variation – has not been quantified (escher et  al. 1975; allaart & jensen 1979). therefore, the subvertical dykes of the main swarm south of itillip ilua have been mapped and their widths measured along six selected transects perpendicular to the dominant dyke orientation in the southern archaean foreland and in the easternmost part of area 5b. figure 3 shows the locations of the six transects (a–f) where the exposure and air photo quality were optimal. results are also summarised in table 1. 3.2.1. inussuttusup tunua (transect a) this area is the same as that covered by the map in escher et  al. (1976, fig. 2), where the focus is the relationship between dykes and shear zones. along this transect (‘a’ in fig. 3), 24 dykes wider than  10  m were recognised across a distance of 5 km. these 24 dykes make up 9.3% of the exposure (table 1). 3.2.2. north of simiutaq (transect b) a total of 17 nne dykes wider than 10 m occur along a distance of 7 km, as measured perpendicular to their area 1 area 12 area 11 ? ?? aujas so q sh ear zone kangerlussuatsiaq maniitsupsermilia 1 32 aujas so q sh ear zone maniitsoq kangaamiut fig. 17 –53° 65°30 66° –52° –51° –50° n 10 km fig. 16 location of the aujassoq shear zone (red dashed lines) relative to the areas and place names cited in the text. 1: maniitsup sermilia fjord. 2: elephant foot valley (informal name). 3: lake tasersiaq. the extent of areas 1, 11 and 12 is also shown. location of map shown in fig. 3. ?: question marks indicate potential continuations of the aujassoq shear zone. 1 km n a b area 11n ew s 20 40 60 80 100 120 140 160 102030% fig. 17 dyke orientations in area 11. a: aerial photo h0945 showing panels of straight, well-foliated and dyke-bearing gneisses centrally in ‘elephant foot valley’. extent of image shown in figs 3 and 16. b: rose diagram for dykes in area 11. location in figs 8 and 16. rose diagram construction described in fig. 9. https://doi.org/10.34194/3cy30452 http://www.geusbulletin.org/ sørensen et al. 2025: geus bulletin 59. 8380. https://doi.org/10.34194/3cy30452 17 of 29 geusbulletin.org strike (fig. 21). in addition, a few ese dykes occur. the nne dykes make up 7.8% of the outcrop (table 1). 3.2.3. tunu (transect c) along a distance of 8.5 km on the north shore of tunu (‘c’ in fig. 3), 31 nne dykes occur, making up 5.7% of the outcrop. additionally, a few ese dykes also occur. the dykes here are slightly more numerous but not as wide as dykes along the coast north of simiutaq (transect b). the dyke densities derived for simiutaq and tunu (table 1) are comparable to the 7% observed by allaart & jensen (1979) and significantly higher than the 2–3% observed by escher et al. (1975). in these profiles, maximum widths are in the range 40–50 m. while seven dykes in the simiutaq transect are 40 m wide, only one 40 m wide dyke occurs within the tunu transect. 3.2.4. kangerlussuaq (transect d) transect d is located 60 km from the mouth of kangerlussuaq. in this area, 10 major dykes occur across a distance of 3 km, with a cumulative width of 315 m, giving a dyke density of 10.5% (table 1). 3.2.5. kangerlussuatsiaq (transect e) approximately 20 km sw of the head of kangerlussuatsiaq, we observe 11 dykes across a distance of 4.1 km, with a cumulative width of 410 m, giving a dyke density of 10.3% (table 1). 3.2.6. inuppaat quuat (transect f) in the area north of the isz, dyke densities have been assessed in area 5b just south of the glacier inuppaat quuat (location in fig. 3; figs 8, 22). the length of the transect is 6 km with a dyke density of 9.9%. almost half of this is due to three exceptionally wide dykes (50, 80 and 120 m) in the western part of the transect (figs 12, 22; table 1). 3.2.7. summary of the 123 dykes tabulated in table 1, only 11 are wider than 40 m, only six wider than 50 m and the widest dyke area 12 n ew s 20 40 60 80 100 120 140 160 102030% fig. 18 rose diagram of dyke orientations from area 12 just south of lake tasersiaq (location in figs 8 and 16). rose diagram construction described in fig. 9. n ew s 20 40 60 80 100 120 140 160 102030% area 10a fig. 19 rose diagram of dyke orientations from area 10a just east of kangaamiut sermiat. location in fig. 8. rose diagram construction described in fig. 9. table 1 dyke characteristics observed along profiles a–f profile profile location details (fig. 3) profile length (km) number of dykes cumulative width (m) % of outcrop a inussuttusup tunua 5 24 465 9.3 b simiutaq (mouth of kangerlussuaq) 7 17 545 7.8 c akuliaruseq 8.5 31 485 5.7 d kangerlussuaq (60 km from mouth) 3 10 315 10.5 e kangerlussuatsiaq (20 km from head) 4 11 410 10.3 f inuppaat quuat 6 30 595 9.9 total: 33.5 123 2815 8.1 average number of dykes per km = 3.7. average width of dykes = 25 m. location and orientation of profiles shown in fig. 3. https://doi.org/10.34194/3cy30452 http://www.geusbulletin.org/ sørensen et al. 2025: geus bulletin 59. 8380. https://doi.org/10.34194/3cy30452 18 of 29 geusbulletin.org fig. 20 the torsuup swarm (informal name) of ese dykes in areas 10b and 10c. a: aerial photo k1477 from the eastern part of area 10b showing how dykes give rise to ese-trending lineaments. n–s width of photo c. 13 km. location in fig 3. b: rose diagrams showing dyke orientations in areas 10b and 10c. locations in fig. 8. rose diagram construction described in fig. 9. 2 km a n b n ew s 20 40 60 80 100 120 140 160 102030% n ew s 20 40 60 80 100 120 140 160 102030% area 10b area 10c https://doi.org/10.34194/3cy30452 http://www.geusbulletin.org/ sørensen et al. 2025: geus bulletin 59. 8380. https://doi.org/10.34194/3cy30452 19 of 29 geusbulletin.org encountered is 120 m. the widest dyke reported in the literature from this area is 140 m wide (mayborn 2000) and occurs in the kangerluarsussuaq fjord c. 20 km n of the mouth of kangerlussuaq (supplementary file s1, figs 3, 7). dykes less than 50 m wide account for 80% of the dyke volume, while the six dykes wider than 50 m make up the remaining 20%. in the foreland region, we have only observed local dyke densities significantly exceeding those listed here adjacent to the nunatak nunakerlua, situated at the northern margin of kangaamiut sermiat south of sarfartoq (‘23’ in fig. 3). in this area, c. 20 dykes with an average width of c. 20 m make up 20% of a c. 0.8 km wide outcrop, part of which is shown in fig. 23. from transects a–e, we conclude that the exposed central part of the dyke swarm in the southern archaean foreland is characterised by densities of 5–10% with a maximum density of c. 20% on and adjacent to nunakerlua. extrapolating the numbers in table 1 (on average 3.7 dykes/km in transects a to e) to the entire swarm (c.  100  km from inussuttusup tunua to the head of kangerlussuatsiaq) would suggest the number of dykes in the central swarm exposed onshore to be c. 370. 2 km n simiutaq kangerlussuaq fig. 21 aerial photo d3208 showing ne-trending dykes for profile ‘b’ (location in fig. 3) along the coast north of simiutaq (lower left) at the mouth of kangerlussuaq. yellow dashed lines indicate the outline of a dyke, c. 80 m wide. 2580 20 50 10 120 30 10 10 20 80 10 1010 10 10 2010 10 10 10 10 10 20 10 631 580 600 539 685 623 49°50'w 49°50'w 49°55'w 49°55'w 50°w 50°w 67 °3 0' n 67 °3 0' n 67 °2 8' n 67 °2 8' n 0 0.5 1 1.5 2 2.5 km inuppaat quuat fig. 22 dyke swarm (green lines) mapped for profile f (location in fig. 3) south of inuppaat quuat. thicknesses are measured in metres indicated for the dykes included in table 1. numbers in italics indicate selected elevations in metres. https://doi.org/10.34194/3cy30452 http://www.geusbulletin.org/ http://maximum sørensen et al. 2025: geus bulletin 59. 8380. https://doi.org/10.34194/3cy30452 20 of 29 geusbulletin.org 3.3. dykes in the isz the isz contains multiple panels of juvenile palaeoproterozoic metasediments in which dykes are absent (e.g. fig. 24). detrital zircons in a sample of these metasediments yielded maximum ages of c. 2100 ma (sample 413840 in nutman et al. 1999). the metasediments are tectonically interleaved with gneisses containing strongly deformed and metamorphosed kangâmiut dykes parallel to the gneiss fabric (fig. 24). in the southern foreland described so far, similar juvenile metasediments have not been identified. 3.3.1. dyke densities in the isz, sarfannguit nunaat we have evaluated dyke densities as close to the coast as exposure permits, near utoqqaat, across sarfannguit nunaat and the westernmost part of sallersua (locations in figs 3, 25). according to grocott (1979), kangâmiut dykes extend northwards near the south coast of maniitsorsuaq, which was suggested to mark the northern limit of the swarm. the southern boundary of the sisimiut charnockite is a several kilometres wide transition zone with panels of charnockite intercalated with kangâmiut dyke bearing gneisses. rocks in the transition zone are strongly deformed and metamorphosed and can be hard to distinguish. high densities of dykes extend for at least 10 km south of this limit. it has been possible to estimate dyke densities in these areas, which are assumed to represent the deformed equivalents of the part of the dyke swarm submerged in the davis strait to the west. from the coast at utoqqaat southwards and across the island of sarfannguit nunaat, high dyke densities occur across a distance of c. 8 km. where densities can be evaluated in this part of the profile, we observe up to 40 dykes per km. this is an order of magnitude larger than the density calculated for the swarm in the foreland while the number of dykes (8 × 40) is comparable to those in the foreland in the 100 km transect from the mouth of itilip ilua to the head of kangerlussuatsiaq (section 3.2). we therefore conclude that dyke densities comparable to that onshore in the foreland are expected to extend offshore. this is likely to be a minimum estimate of the dyke density since some thin dykes identified in the foreland may not be visible on the imagery used in the utoqqaat to sallersua area. 3.4. dykes in the transition zone in the area se of the isz, deformation is highly variable. here, we term this the ‘transition zone’ (fig. 7). 300 m n 300 m fig. 23 ‘extreme’ dyke density of c. 20% is visible on (part of) the nunataq nunakerlua in the northern border zone of kangaamiut sermiat. location in fig. 3. source: google earth. fig. 24 panel of rusty weathering metasediments (graphiteand sulphide-bearing) interleaved with gneisses containing deformed kangâmiut dykes parallel with the nagssugtoqidian fabric of the gneisses. view towards the east, height of cliff c. 400 m, full width of exposure 1.3 km. location in fig. 3. https://doi.org/10.34194/3cy30452 http://www.geusbulletin.org/ sørensen et al. 2025: geus bulletin 59. 8380. https://doi.org/10.34194/3cy30452 21 of 29 geusbulletin.org fig. 25 dykes on the island of sarfannguit nunaat, se of sisimiut as seen in aerial photo e3128. location in fig. 3. a: zoomed out view. area covered in panel b marked in yellow. b: high dyke density is visible in the central part of sarfannguit nunaat (width of image 7 km). 1 km n a b b 8 km n sarfannguit nunaat maniitsorsuaq utoqqat sallersua https://doi.org/10.34194/3cy30452 http://www.geusbulletin.org/ sørensen et al. 2025: geus bulletin 59. 8380. https://doi.org/10.34194/3cy30452 22 of 29 geusbulletin.org it is the area between the isz, the snf (fig. 2) and itillip ilua fjord. this transition zone encompasses parts of areas 4 and 6, all of area 7 and parts of areas 8 and 9 (fig. 8). folding is conspicuous in parts of the transition zone. the structural changes from the isz and the area to the south are illustrated in fig.  26, demonstrating the contrast between isz (inner maligiaq, figs 3, 26a) and the transition zone (kangerluarsuk; saqqap kangerluarsua; figs  3, 26b) as mapped by korstgård (1980). 3.4.1. areas 4, 6 and 7 north of itillip ilua, ene dykes are dominant and likely to be reorientated nne dykes (fig. 27, area 4). however, ese dykes are also present. dykes from area 1 to area 4 thus represent a transition from the original nne orientation in the foreland to an ene-trending nagssugtoqidian orientation. this transition occurs across a narrow zone extending eastwards along itillip ilua, across kangerlussuaq and along sarfartoq valley (fig. 5). the relationship between dyke area 7 n ew s 20 40 60 80 100 120 140 160 102030% area 6 n ew s 20 40 60 80 100 120 140 160 102030% area 4 n ew s 20 40 60 80 100 120 140 160 102030% fig. 27 rose diagrams showing typical nagssugtoqidian orientations of dykes in areas 4, 6 and 7 representing the zone south of the corridor of highly parallelised dykes and gneiss structures in the ikertooq shear zone and north of the southern nagssugtoqidian front. the extent of these areas is shown in fig. 8. rose diagram construction described in fig. 9. n = 375 n = 197 n = 34 n = 101 a binner maligiaq saqqap kangerluarsua fig. 26 stereographic plots (lower hemisphere projection) show both the poles of planar (blue) and linear (red) structures in the saqqap kangerluarsua – ikertooq region. structural orientation data covers the ikertooq shear zone (isz) and the transition zone to the south. a: data from the inner part of ikertooq (inner maligiaq, ‘26a’ in fig. 3). b: data from the transition zone to the south of the isz (saqqap kangerluarsua, ‘26b’ in fig. 3). the complete girdle in b reflects the pervasive and consistent folding in the transition zone. structural measurements are further contoured, using stereonet v. 11 software (allmendinger et al. 2012; cardozo & allmendinger 2013). kamb contouring, interval 2 sigma. data from korstgård (1980). https://doi.org/10.34194/3cy30452 http://www.geusbulletin.org/ sørensen et al. 2025: geus bulletin 59. 8380. https://doi.org/10.34194/3cy30452 23 of 29 geusbulletin.org orientation and lineament geometry is discussed in detail elsewhere (korstgård et al. 2024). the nagssugtoqidian trend observed towards the east in area 4 continues into areas 6 and 7 (figs 8, 27), thus collectively outlining a zone of predominant nagssugtoqidian trends between the southern foreland and the northern corridor of highly parallelised dykes and gneiss structures in the isz. fig. 28 rose diagrams of dyke orientations for areas 8 and 9 in the foreland immediately south of the southern nagssugtoqidian front. the extent of these areas is shown in fig. 8. rose diagram construction described in fig. 9. n ew s 20 40 60 80 100 120 140 160 102030% area 8 n ew s 20 40 60 80 100 120 140 160 102030% area 9 fig. 29 extremely deformed dykes are visible in the northern wall (height of section c. 700 m) of sarfartoq valley (‘escher’s wall’). this exposure is part of the paradise thrust segment of hageskov (1995b) and olsen (2000). in this wall, c. 50 dykes are visible, and detailed measurements indicate they make up 22% of the exposure. location in fig. 3. valley trends e–w, viewing direction towards the ne. https://doi.org/10.34194/3cy30452 http://www.geusbulletin.org/ sørensen et al. 2025: geus bulletin 59. 8380. https://doi.org/10.34194/3cy30452 24 of 29 geusbulletin.org 3.4.2. areas 8 and 9 in areas 8 and 9 (figs 8, 28), nagssugtoqidian deformation is less intense than in the adjacent area 7 where zones of stronger deformation are prevalent. on the original escher et al. (1975) map, shown here as fig. 5, the nne dykes in areas 8 and 9 were marked as undeformed. however from our mapping it appears that the nne and ese dykes have been significantly deformed (see figs 7, 28). the boundary between areas 6 and 7 and areas 8 and 9 corresponds to the “southern nagssugtoqidian boundary” or “main nagssugtoqidian boundary” of escher et al. (1975, 1976), “southern nagssugtoqidian structural front” of hageskov (1995a) or our preferred “southern nagssugtoqidian front” (snf) of van gool et al. (2002). escher et al.’s (1975) illustration of this boundary (their figs 8 & 10) features the north side of sarfartoq valley (fig. 29) showing the strongly parallelised dykes and host rock structures in what hageskov (1995a, 1995b) calls part of a ‘paradise thrust segment’. as a participant in the dlc project, olsen (2000) mapped in the area from sarfartoq to c. 10 km east of the snf. this mapping included the complex geometry of the snf and added some detail to the southern part of hageskov’s generalised segmentation model and is shown in fig. 5. olsen (2000) noted the marked increase in deformation from kangerlussuaq south-eastwards in sarfartoq valley. the presence of undeformed nne-trending and e–w-trending kangâmiut dykes with metamorphic mineral assemblages was also observed. it is clear that structural heterogeneity characterises this zone and that it is a distinctive feature of the entire snf with weakly deformed dykes interspersed with zones of intense deformation. 3.4.3. the transition zone near the head of kangerlussuaq a spectacular example of the structural transition across the snf is visible to those entering greenland via kangerlussuaq airport. an area near the airport has been described by engström & klint (2014), who report two generations of post-dyke ductile shear zones, the oldest of which is dextral. one of these is shown here (fig. 30a). the density of dykes in this shear zone is comparable to the density of dykes in the isz described earlier. this is in contrast to the area se of the runway where weakly deformed ese dykes can be seen (fig. 30b). these features are also evident in supplementary file s1 and in figs 6b & 6c of engström & klint (2014). the areas described by engström & klint (2014) near the head of kangerlussuaq are situated within the transition zone. they perfectly summarise the variability of structures in this zone (p. 326) as a change in deformation style of the kangâmiut dykes, which they observed along a regional n–s transect from a “folded area in the north” and through “the strongly sheared system 1 shear zones in the central area, towards the generally undisturbed area south of the sandflugtdalen thrust fault.” for reference, the ‘system 1 shear zones’ are e–w-trending and an example is shown in fig. 30a. the ‘sandflugtdalen thrust fault’ trends ene–wsw and is located between the areas shown in figs. 30a and b. 4. discussion the main nagssugtoqidian structural features and a schematic representation of the kangâmiut dykes interpreted in this study are shown in fig. 7 (and in full in supplementary file s1). a significant new observation resulting from our mapping of these dykes is that mafic dykes – specifically members of the kangâmiut dyke swarm – may occur north of the isz. this interpretation is open to conjecture. however, considering the potential implications, we will discuss possible consequences for the nature of the isz and for the evolution of the sno. if future work shows that alternative interpretations are more likely, the tectonic implications discussed here will need to be adjusted. n 100 m a b fig. 30 dykes visible near kangerlussuaq airport. a: strongly deformed dykes north of the runway. view towards ne, exposure is c. 525 m long. b: weakly deformed subvertical 15–20 m wide e–w trending dykes sse of runway (similar to the relationships in fig. 6b of engström & klint 2004). major dykes indicated by arrows. locations in fig. 3. source: google earth. https://doi.org/10.34194/3cy30452 http://www.geusbulletin.org/ sørensen et al. 2025: geus bulletin 59. 8380. https://doi.org/10.34194/3cy30452 25 of 29 geusbulletin.org fig. 31 proposed model to explain the distribution of kangâmiut dykes in the central and southern nagssugtoqidian orogens (cno, sno). x marks the proposed westernmost occurrence of kangâmiut dykes south of ikertooq shear zone (isz). y and z mark the westernmost extent of kangâmiut dykes north of the isz, depending on whether this margin is interpreted to be 100 km (y) or 30 km (z) west of the greenland ice sheet. if the dykes at y were once co-extensive with the western margin of the swarm south of the isz, they would have originated near y* assuming translation perpendicular to the isz. similarly, if the dykes at z represented the western limit of the swarm north of the isz, they would necessarily have come from near z*. kangâmiut-1: exploration well. solid dykes are generalised from observations shown in fig. 7. stippled dykes are extrapolated. blue dykes are from in or near to the isz and their extrapolated equivalents. red dykes are those observed south of the isz and their extrapolated equivalents. grey arrows: direction of shortening in the two scenarios described in the text (y*→y and z*→z). 50 km schematic distribution of the kangâmiut ma�c dyke swarm: observed interpreted 67° 68°68° 67° 52°54° 50° 48° 69° 66° 54° 52° n nordre isortoq shear zone ikertooq thrust zone s n f sisimiut kangaamiut sermiat s n f greenland ice sheet attu qeqertarsuup tunua qasigiannguit ilulissat aasiaat kangaamiut kangâmiut-1 maniitsoq nordre strøm�ord shear zone nordre isortoq shear zone ikertooq shear zone c n o n n o sn o iso rto q kangerlussu atsia q kanger lus suaq x y z y* z* https://doi.org/10.34194/3cy30452 http://www.geusbulletin.org/ sørensen et al. 2025: geus bulletin 59. 8380. https://doi.org/10.34194/3cy30452 26 of 29 geusbulletin.org in the isz, the intercalation of dyke-bearing gneiss panels with juvenile sediments (see fig. 24) leaves no doubt that this structure is of prime tectonic significance. it has been interpreted as a ductile zone of sto sse-directed thrusting (korstgård et  al. 1987). the isz separates areas characterised by distinctly different distributions of kangâmiut dykes. in the zone between the isz and the nordre isortoq shear zone (fig. 2), the approximate western limit of the dense part of the nne swarm is suggested (section 3.1.3) to be located either c. 100 km west of the greenland ice sheet (including the area shown in fig. 13) or to the west of the inuppaat quuat area (fig. 12a) c. 30 km west of the greenland ice sheet. the eastern limit is presumably covered by the greenland ice sheet. this contrasts with the area south of the isz where an approximate eastern limit is visible, but a western limit is not exposed. south of itillip ilua (fig. 7), a high density of dykes is seen southwards to maniitsoq. the eastern limit of this high-density core of the swarm can be traced towards the nne and ne in the vicinity of the aujassoq shear zone. as discussed in section 3.3, the western limit to the swarm must lie c. 100 km offshore if the density of dykes offshore is comparable to that onshore. that the kangâmiut-1 well is situated 100 km offshore may be seen as evidence that this suggested extent of the swarm offshore is not unreasonable. it is our hypothesis that the dykes north and south of the isz constituted a continuous swarm prior to nagssugtoqidian tectonism, and that crustal shortening across the isz brought a northern and southern segment into their present relative positions. to assess the amount of shortening across the isz, some simplifications are needed. the isz is reduced to a discrete line, x–y–z in fig. 31 with ‘x’ representing the western limit of the swarm south of the isz. line x–y*–z* represents the western limit of the swarm prior to thrusting (fig. 31). the x–y*–z* trend is parallel to the average trend of kangâmiut dykes in the southern archaean foreland. north of the isz, the western limit of the dykes was suggested to be either c. 30 km or c. 100 km west of inuppaat quuat (section 3.1.3). in the former case, the eastward shift of the swarm is c. 230 km (x–z in fig. 31), while in the latter scenario (x–y), the shift would be c. 150 km. to explain these shifts as a result of southward thrusting in a direction perpendicular to the isz (as illustrated by grey arrows in fig. 31), we arrive at distances between 160 km (y*–y in fig. 31) and 260 km (z*–z), depending on how the western extent of the swarm north of isz is interpreted. it should be noted that this magnitude of thrusting agrees with the extent of crustal loading southwards from ikertooq inferred independently by korstgård et al. (2024) from metamorphic overprinting of the kangâmiut dykes. one of the consequences of this model is that the crustal segments now juxtaposed across the isz (cno and sno, fig. 2) at the time of dyke emplacement were separated by 160–260 km of intervening, dyke-bearing continental crust. if the nordre strømfjord shear zone and nordre isortoq shear zone likewise represent the loci of significant crustal shortening, the present-day width of the cno (c. 100 km from the northern margin of the nordre strømfjord shear zone to the southern margin of the isz) represents only a small fraction of the pre-nagssugtoqidian continental crust in west greenland. 5. conclusions the kangâmiut dyke swarm has in many studies been shown to be a sensitive monitor of the structural and metamorphic changes occurring across the southern margin of the nagssugtoqidian orogen. with that in mind, the present effort set out to quantify the change in orientation and density of dykes across the extent of the swarm. this was accomplished via publicly available aerial imagery and involved careful visual inspection of c. 125 aerial photos in 1:150 000 scale, locally complemented with eox and google earth satellite images. the snf is marked by a zone of south-directed thrusts in which dykes and their host gneisses are strongly parallelised. south of the snf – in the foreland – the swarm is nne-trending and contains a minor subset of dykes with other orientations, dominantly e–w to ese–wnw. from the south towards the snf, the orientation of the dykes gradually rotates from nne to the characteristic ene-orientated nagssugtoqidian trend. notably, nne dykes are rotated clockwise, whereas e and ese dykes are rotated counterclockwise, thus indicating an overall nnw-orientated shortening direction. in several profiles both south and north of the snf, we made detailed measurements of dyke orientation, density and thickness. in these profiles, dykes were found to make up 6–11% of the outcrop, thus indicating the magnitude of crustal extension accompanying dyke emplacement. of the 123 dykes observed, the average dyke width is 25 m, with 80% of these being less than 50 m wide. while the kangâmiut dyke swarm is generally believed to not extend northwards beyond the isz, the aerial images studied here of the eastern area between the isz and the nordre isortoq shear zone revealed what appears to be a group of nneto ne-trending dykes. their belonging to the kangâmiut dyke swarm is at this point speculative, however should this be confirmed, it has interesting implications for the tectonic evolution of the isz. the western extent of the dyke swarm south of the isz is not exposed and must be somewhere in the https://doi.org/10.34194/3cy30452 http://www.geusbulletin.org/ sørensen et al. 2025: geus bulletin 59. 8380. https://doi.org/10.34194/3cy30452 27 of 29 geusbulletin.org davis strait. if our assumptions about the extent of the kangâmiut dyke swarm are correct, the calculated shortening across the isz is in the order of 150–300 km, which is the same order of magnitude proposed by recent studies based on metamorphic overprints and geochronological data from the southern nagssugtoqidian orogen and its foreland. acknowledgements the authors extend their appreciation to jeroen van gool and kyle mayborn for careful, critical and constructive reviews of earlier versions of this manuscript. it is also their pleasure to acknowledge geus’ editorial staff for their patience and help. additional information funding statement this study received no funding. competing interests the authors declare no competing interests. author contributions ks, jk: conceptualization, investigation, writing. fcm: writing, editing. ww: maps. additional files supplementary file s1 contains the revised map of kangâmiut dykes and major structures of the nagssugtoqidian orogen and is available at https://doi.org/10.22008/fk2/somuf1 references allaart, j.h. 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(eds): early precambrian processes. geological society of london, special publications 95, 211–224. https://doi.org/10.1144/gsl. sp.1995.095.01.13 park, g. 2022: a regional explanation for laxfordian tectonic evolution and its implications for the lewisian terrane model. scottish journal of geology 58(1). https://doi.org/10.1144/sjg2021-020 pedersen, m., weng, w.l., keulen, n. & kokfelt, t.f. 2013: a new seamless digital 1:500 000 scale geological map of greenland. geological survey of denmark and greenland bulletin 28, 65–68. https://doi. org/10.34194/geusb.v28.4727 ramberg, h. 1949: on the petrogenesis of the gneiss complexes between sukkertoppen and christanshaab, west greenland. meddelelser fra dansk geologisk forening 11, 312–327. sørensen, k. 1970: some observations on the structural and metamorphic chronology on agto and surrounding islands, central west greenland. rapport grønlands geologiske undersøgelse 27, 32 pp. https://doi.org/10.34194/rapggu.v27.7219 sørensen, k. 1971: amfibolitfacies – granulitfacies overgangen i metadoleritter fra agtoområdet, vestgrønland. 65 pp. msc thesis, department of geology, aarhus university. sørensen, k. 1977: report on the rock types between levels 3672 and 3874 at the kangâmiut 1 site. local setting and regional context. 63 pp. total grønland olie a/s paris. sørensen, k., korstgård, j.a., glassley, w.e. & stensgaard, b.m. 2006: the nordre strømfjord shear zone and the arfersiorfik quartz diorite in arfersiorfik, the nagssugtoqidian orogen, west greenland. geological survey of denmark and greenland bulletin 11, 145–161. https://doi. org/10.34194/geusb.v11.4922 st-onge, m.r., van gool, j.a.m., garde, a.a. & scott, d.j. 2009: correlation of archaean and palaeoproterozoic units between northeastern canada and western greenland: constraining the pre-collisional upper plate accretionary history of the trans-hudson orogen. geological society, london, special publications 318, 193–235. https://doi. org/10.1144/sp318.7 van gool, j.a.m., connelly, j.n., marker, m. & mengel, f.c. 2002: the nagssugtoqidian orogen of west greenland: tectonic evolution and regional correlations from a west greenland perspective. canadian journal of earth sciences 39, 665–686. https://doi.org/10.1139/ e02-027 van gool, j.a.m. & marker, m. (compilers) 2004: geological map of greenland, 1:100  000, ussuit, 67 v.2 nord. copenhagen: geological survey of denmark and greenland. https://doi.org/10.22008/ fk2/3vzspo van gool, j.a.m. & marker, m. 2007: explanatory notes to the geological map of greenland, 1:100  000, ussuit 67 v.2 nord. geological survey of denmark and greenland map series 3, 40 pp. https://doi. org/10.34194/geusm.v3.4596 watterson, j. 1974: investigations of the nagssugtoqidian boundary in the holsteinsborg district, central west greenland. rapport grønlands geologiske undersøgelse 65, 33–37. https://doi.org/10.34194/ rapggu.v65.7383 windley, b.f. 1964–1967: field map of area around manitsup sermilia. map id 1589 (1:50 000). copenhagen: geological survey of denmark and greenland. windley, b.f. 1970: primary quartz ferro-dolerite/garnet amphibolite dykes in the sukkertoppen region of west greenland. in: newall, g. & rast, n. (eds): mechanism of igneous intrusion. geological journal special issue 2, 79–92. zeck, h.p. & kalsbeek, f. 1981: geochemistry of amphibolite facies metamorphism of a suite of basic dykes, precambrian basement, greenland. chemie der erde 40, 1–22. https://doi.org/10.34194/3cy30452 http://www.geusbulletin.org/ https://doi.org/10.34194/rapggu.v89.7565 https://doi.org/10.1016/j.precamres.2018.12.028 https://doi.org/10.1016/j.precamres.2018.12.028 https://doi.org/10.22008/fk2/oc5wns https://doi.org/10.22008/fk2/oc5wns https://doi.org/10.2475/04.2008.06 https://doi.org/10.1016/s0301-9268(98)00099-0 https://doi.org/10.1016/s0301-9268(98)00099-0 https://doi.org/10.22008/fk2/ecxcql https://doi.org/10.22008/fk2/ecxcql https://doi.org/10.34194/rapggu.v89.7563 https://doi.org/10.34194/rapggu.v89.7563 https://doi.org/10.1144/gsl.sp.1995.095.01.13 https://doi.org/10.1144/gsl.sp.1995.095.01.13 https://doi.org/10.1144/sjg2021-020 https://doi.org/10.34194/geusb.v28.4727 https://doi.org/10.34194/geusb.v28.4727 https://doi.org/10.34194/rapggu.v27.7219 https://doi.org/10.34194/geusb.v11.4922 https://doi.org/10.34194/geusb.v11.4922 https://doi.org/10.1144/sp318.7 https://doi.org/10.1144/sp318.7 https://doi.org/10.1139/e02-027 https://doi.org/10.1139/e02-027 https://doi.org/10.22008/fk2/3vzspo https://doi.org/10.22008/fk2/3vzspo https://doi.org/10.34194/geusm.v3.4596 https://doi.org/10.34194/geusm.v3.4596 https://doi.org/10.34194/rapggu.v65.7383 https://doi.org/10.34194/rapggu.v65.7383 the kangâmiut dyke swarm in west greenland: a new map and insights into their tectonic evolution 1. introduction 1.1. early work on the nagssugtoqidian orogen 1.2. previous research on the kangâmiut dykes 1.3. regional geological framework 2. new mapping 2.1. data 2.2. mapping objectives 3. description of the dyke map 3.1. dykes in the southern foreland and area 5 3.1.1. areas 1-4 3.1.2. age relations between nne and ese dykes 3.1.3. area 5 3.1.4. areas 11, 12 and the aujassoq shear zone 3.1.5. area 10a 3.1.6. areas 10b and 10c 3.2. dyke densities in the southern archaean foreland and in area 5 3.2.1. inussuttusup tunua (transect a) 3.2.2. north of simiutaq (transect b) 3.2.3. tunu (transect c) 3.2.4. kangerlussuaq (transect d) 3.2.5. kangerlussuatsiaq (transect e) 3.2.6. inuppaat quuat (transect f) 3.2.7. summary 3.3. dykes in the isz 3.3.1. dyke densities in the isz, sarfannguit nunaat 3.4. dykes in the transition zone 3.4.1. areas 4, 6 and 7 3.4.2. areas 8 and 9 3.4.3. the transition zone near the head of kangerlussuaq 4. discussion 5. conclusions acknowledgements additional information funding statement competing interests author contributions additional files references figures fig. 1 original overview map of the nagssugtoqidian from ramberg (1949), with their place name spellings, showing the proposed subdivision of the precambrian rocks between kangaamiut (kangâmiut in the figure) and qasigiannguit (christianshaab in the figure) in west greenland. the isortoq gneiss complex is in granulite facies, while the adjoining ikertoq and egedesminde gneiss complexes are in amphibolite facies. slightly modified from ramberg (1949). fig. 2 simplified geological map of west greenland, showing the internal subdivisions of the nagssugtoqidian orogen (sno: southern nagssugtoqidian orogen, cno: central nagssugtoqidian orogen, nno: northern nagssugtoqidian orogen, after marker et al. 1995). the kangâmiut dykes investigated in this study occur from 65°10´n to 67°45´n. north of itillip ilua and the southern nagssugtoqidian front (snf), country rocks are largely archaean orthogneisses and minor supracrustal rocks, which are all variably to completely reworked during nagssugtoqidian orogenesis. the thin stippled line at kangaamiut sermiat traces the aujassoq shear zone (“aujassoq-evighedsfjord shear belt” of allaart & jensen 1978). ‘kangâmiut-1’ is the location of the well drilled in 1976 by the french oil company total. black triangles: thrust faults. inset map shows the extent of the nagssugtoqidian orogen across southern greenland. slightly modified from korstgård et al. (2024) and includes information from escher (1971) and allaart (1982). fig. 3 index map showing the locations of figures, photos and place names cited in the text. extents of aerial photos and maps are shown with the prefix ‘fig’. location of field photographs and other displays are shown by the figure number. red lines: trends and lengths of transects a–f are shown schematically. grey stippled lines: simplified major tectonic boundaries. fig. 4 original reconnaissance map from noe-nygaard & ramberg (1961), with their place name spellings. location in fig. 3. diagonal shading with diamond symbols: agmatitic hypersthene-gneiss and pyroxene-gneiss of enderbitic to quartz-dioritic composition. black lines: observed kangâmiut dykes. purple shading: ultrabasic rocks (lower right of map). yellow shading: granulite facies metasediments (lower right of map). slightly modified from noe-nygaard & ramberg (1961). fig. 5 original map of kangâmiut dykes from escher et al. (1975). location in fig. 3. the frame in the sw corner shows the area covered by fig. 4. snf: southern nagssugtoqidian front. the trace of the snf from escher et al. (1975) and mapping by hageskov (1995a, 1995b) and olsen (2000). modified from korstgård et al. (2024). fig. 6 aerial photographs of kangâmiut dykes. locations in fig. 3. a: aerial photo f1946. kangâmiut dykes hosted by amphibolite facies gneisses in kangerlussuaq c. 57 km from the mouth of the fjord. width of the fjord sw of the delta lobe is 2.4 km. the frame indicates the extent of the outcrop shown in fig. 11a. b: aerial photo e3116. kangâmiut dykes in granulite facies host gneiss showing a clear intersection between nne and a later ese dyke, north shore of kangerlussuaq, 20 km from the mouth of the fjord. the ese-trending dyke is 80 m wide. fig. 7 major nagssugtoqidian structures and generalised trends of dykes based on interpretation of aerial photos. this map is a simplified version of supplementary file s1. red lines: structural trends in the major shear and thrust zones. black lines: general trend of dykes where observable as continuous, fig. 8 map showing the extent of areas 1–12 with rose diagrams of dyke orientations for each area. grey stippled lines: simplified major tectonic boundaries. nssz: nordre strømfjord shear zone, nisz: nordre isortoq shear zone, isz: ikertooq shear zone, snf: southern nagssugtoqidian front. fig. 9 rose diagrams for areas 1, 2, 3 and 4 demonstrating gradual clockwise change in orientation of the nne dykes towards the ‘typical’ nagssugtoqidian ene trend. location of areas 1–4 shown in fig. 8. within each area, azimuths of dykes are grouped in 10-degree bins and lengths of dykes are summed and represented as percentages of cumulative length. fig. 10 field photographs showing the three stages – a to b to c – in the progressive deformation of the kangâmiut dykes as seen in outcrops along the south shore of kangerlussuaq. locations of photos in fig. 3. a: the arrow points to the peak named ‘stoppenålen’, the nw face of which (left side in image) is formed by a nne dyke dipping c. 80° nnw. viewing direction is ene. b: dykes strike ne and dip c. 50 degrees to nw. viewing fig. 11 examples of mutually intersecting dykes. a: ese-dyke cutting (at ‘1’) and being cut (at ‘2’) by nne dykes. cliff face is c. 1.3 km wide. width of the widest dyke c. 50 m. location in figs 3, 6a. source: aerial photo f1946. b: ese dyke (dyke 2) cutting as well as being cut by nne dykes (dyke 1 and dyke 3). image is 1 km wide. dykes are 50–60 m wide. location in fig. 3. source: google earth. fig. 12 dykes in the eastern part of area 5b (location in fig. 8). these dykes are informally named the inuppaat swarm after the inuppaat quuat glacier visible in the figure. a: overview of the area. stippled white line is a major lineament referred to in the text. eox image. box shows the extent of fig. 12b. b: aerial photo l0840 showing an enlarged view of several large dykes of the inuppaat swarm. dykes that are readily visible at this scale are indicated by arrows. widest dyke is 120 m. fig. 13 n–nne-trending dykes south of the meltwater plain isortup kuua. location in fig. 3. e–w width of image is 23 km. the nature of exposure does not permit an assessment of dyke density. only in a few places can a clear colour contrast provide confirmation of a dyke, but terrain lineaments nevertheless suggest the occurrence of a substantial number of dykes. source: aerial photo h927. fig. 14 rose diagrams showing dyke orientations in areas 5a and 5b. locations in fig. 8. rose diagram construction described in fig. 9. fig. 15 nne–ne-trending dykes immediately north of the eastern extension of the ikertooq shear zone (isz). limited colour contrast between dykes and gneisses suggests that host rocks are in granulite facies. a few dykes are indicated with arrows. the southern part of the image shows the e–w trend of the isz. location in fig. 3. source: aerial photo k0349. fig. 16 location of the aujassoq shear zone (red dashed lines) relative to the areas and place names cited in the text. 1: maniitsup sermilia fjord. 2: elephant foot valley (informal name). 3: lake tasersiaq. the extent of areas 1, 11 and 12 is also shown. location of map shown in fig. 3. ?: question marks indicate potential continuations of the aujassoq shear zone. fig. 17 dyke orientations in area 11. a: aerial photo h0945 showing panels of straight, well-foliated and dyke-bearing gneisses centrally in ‘elephant foot valley’. extent of image shown in figs 3 and 16. b: rose diagram for dykes in area 11. location in figs 8 and 16. rose diagram construction described in fig. 9. fig. 18 rose diagram of dyke orientations from area 12 just south of lake tasersiaq (location in figs 8 and 16). rose diagram construction described in fig. 9. fig. 19 rose diagram of dyke orientations from area 10a just east of kangaamiut sermiat. location in fig. 8. rose diagram construction described in fig. 9. fig. 20 the torsuup swarm (informal name) of ese dykes in areas 10b and 10c. a: aerial photo k1477 from the eastern part of area 10b showing how dykes give rise to ese-trending lineaments. n–s width of photo c. 13 km. location in fig 3. b: rose diagrams showing dyke orientations in areas 10b and 10c. locations in fig. 8. rose diagram construction described in fig. 9. fig. 21 aerial photo d3208 showing ne-trending dykes for profile ‘b’ (location in fig. 3) along the coast north of simiutaq (lower left) at the mouth of kangerlussuaq. yellow dashed lines indicate the outline of a dyke, c. 80 m wide. fig. 22 dyke swarm (green lines) mapped for profile f (location in fig. 3) south of inuppaat quuat. thicknesses are measured in metres indicated for the dykes included in table 1. numbers in italics indicate selected elevations in metres. fig. 23 ‘extreme’ dyke density of c. 20% is visible on (part of) the nunataq nunakerlua in the northern border zone of kangaamiut sermiat. location in fig. 3. source: google earth. fig. 24 panel of rusty weathering metasediments (graphiteand sulphide-bearing) interleaved with gneisses containing deformed kangâmiut dykes parallel with the nagssugtoqidian fabric of the gneisses. view towards the east, height of cliff c. 400 m, full width of exposure 1.3 km. location in fig. 3. fig. 25 dykes on the island of sarfannguit nunaat, se of sisimiut as seen in aerial photo e3128. location in fig. 3. a: zoomed out view. area covered in panel b marked in yellow. b: high dyke density is visible in the central part of sarfannguit nunaat (width of image 7 km). fig. 26 stereographic plots (lower hemisphere projection) show both the poles of planar (blue) and linear (red) structures in the saqqap kangerluarsua – ikertooq region. structural orientation data covers the ikertooq shear zone (isz) and the transition zone to the south. a: data from the inner part of ikertooq (inner maligiaq, ‘26a’ in fig. 3). b: data from the transition zone to the south of the isz (saqqap kangerluarsua, ‘26b’ in fig. 3). the complete girdle in b reflects the pervasive and consistent folding in the transition zone. structural measurements are further contoured, using stereonet v. 11 software (allmendinger et al. 2012; cardozo & allmendinger 2013). kamb contouring, interval 2 sigma. data from korstgård (1980). fig. 27 rose diagrams showing typical nagssugtoqidian orientations of dykes in areas 4, 6 and 7 representing the zone south of the corridor of highly parallelised dykes and gneiss structures in the ikertooq shear zone and north of the southern nagssugtoqidian front. the extent of these areas is shown in fig. 8. rose diagram construction described in fig. 9. fig. 28 rose diagrams of dyke orientations for areas 8 and 9 in the foreland immediately south of the southern nagssugtoqidian front. the extent of these areas is shown in fig. 8. rose diagram construction described in fig. 9. fig. 29 extremely deformed dykes are visible in the northern wall (height of section c. 700 m) of sarfartoq valley (‘escher’s wall’). this exposure is part of the paradise thrust segment of hageskov (1995b) and olsen (2000). in this wall, c. 50 dykes are visible, and detailed measurements indicate they make up 22% of the exposure. location in fig. 3. valley trends e–w, viewing direction towards the ne. fig. 30 dykes visible near kangerlussuaq airport. a: strongly deformed dykes north of the runway. view towards ne, exposure is c. 525 m long. b: weakly deformed subvertical 15–20 m wide e–w trending dykes sse of runway (similar to the relationships in fig. 6b of engström & klint 2004). major dykes indicated by arrows. locations in fig. 3. source: google earth. fig. 31 proposed model to explain the distribution of kangâmiut dykes in the central and southern nagssugtoqidian orogens (cno, sno). x marks the proposed westernmost occurrence of kangâmiut dykes south of ikertooq shear zone (isz). y and z mark the westernmost extent of kangâmiut dykes north of the isz, depending on whether this margin is interpreted to be 100 km (y) or 30 km (z) west of the greenland ice sheet. if the dykes at y were once co-extensive with the western margin of the swarm south of the isz, they would have originated near y* assuming translation perpendicular table table 1 dyke characteristics observed along profiles a-f geological survey of denmark and greenland bulletin 3, 1-196 1 geological survey of denmark and greenland bulletin 3 · 2004 late quaternary environmental changes recorded in the danish marine molluscan faunas kaj strand petersen geological survey of denmark and greenland ministry of the environment geus bulletin no 3.pmd 28-06-2004, 08:451 2 geological survey of denmark and greenland bulletin 3 keywords bottom-communities, climate changes, danish, environment, interglacial–glacial cycle, late quaternary, marine, mollusc faunas. cover donax vittatus on the sandy shores of northern france. kaj strand peteresen danmarks og grønlands geologiske undersøgelse øster voldgade 10, dk-1350 copenhagen k, denmark e-mail: ksp@geus.dk scientific editor of this volume: svend stouge editorial secretaries: esben w. glendal and birgit eriksen referees: svend funder and gotfred høpner petersen, denmark illustrations: gurli e. hansen bengaard and henrik klinge pedersen digital photographic work: benny m. schark and jakob lautrup graphic production: knud gr@phic consult, odense, denmark printers: schultz grafisk, albertslund, denmark manuscript submitted: 9 january 1998 final version approved: 11 december 2003 printed: 15 july 2004 isbn 87-7871-122-1 geological survey of denmark and greenland bulletin the series geological survey of denmark and greenland bulletin replaces geology of denmark survey bulletin and geology of greenland survey bulletin. citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 3, 268 pp. available from geological survey of denmark and greenland øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk or geografforlaget aps rugårdsvej 55, dk-5000 odense c, denmark phone: +45 63 44 16 83, fax: +45 63 44 16 97, e-mail: go@geografforlaget.dk © danmarks og grønlands geologiske undersøgelse (geus), 2004 geus bulletin no 3.pmd 28-06-2004, 08:452 3 contents abstract . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7 introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9 danish sites with marine sediments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9 the late pleistocene . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10 the holocene . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 12 the recent fauna of shell-bearing molluscs compared to the subfossil fauna . . . . . . . . . . . 12 molluscan finds within the seven regions during the holocene . . . . . . . . . . . . . . . . . . . . . 15 the bælt sea area . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 15 the baltic area . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16 the kattegat area with fjords . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16 the limfjord area . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17 the north sea . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19 the vendsyssel area . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 20 the skagen well area . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 22 the danish late quaternary marine molluscs . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 24 class polyplacophora . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25 order neoloricata . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25 class gastropoda . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25 subclass prosobranchia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25 order archaeogastropoda . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25 order mesogastropoda . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 28 order heterogastropoda . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 39 order neogastropoda . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 44 subclass heterobranchia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 50 order heterostropha . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 50 subclass opisthobranchia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 55 order bullomorpha . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 55 order anaspidea . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 58 order thecosomata . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 59 order gymnosomata . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 60 subclass pulmonata . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 60 order basommatophora . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 60 class scaphopoda . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 60 order siphonodentalioida . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 60 order dentalioida . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 61 class bivalvia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 62 subclass palaeotaxodonta . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 62 order nuculoida . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 62 subclass pteriomorphia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 66 order arcoida . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 66 order mytiloida . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 67 order pteroida . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 69 subclass heterodonta . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 73 order veneroida . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 73 order myoida . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 92 subclass anomalodesmata . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 95 order pholadomyoida . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 95 geus bulletin no 3.pmd 28-06-2004, 08:453 4 the skagen well . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 98 the skagen well – perspectives . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 98 the pre-late quaternary deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 99 the late pleistocene . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 99 eemian deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 99 the early/middle weichselian, marine and glacigene deposits . . . . . . . . . . . . . . . . . . 100 the late weichselian marine and glacigene deposits . . . . . . . . . . . . . . . . . . . . . . . . . 101 the holocene . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 102 the preboreal–boreal 10 000 – 8000 14c years b.p. . . . . . . . . . . . . . . . . . . . . . . . . . . 103 the atlantic 8000–5000 14c years b.p. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 104 the subboreal 5000–2500 14c years b.p. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 105 the subatlantic 2500– 14c years b.p. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 105 the older subatlantic . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 106 the younger subatlantic . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 107 conclusive remarks on the skagen well . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 112 the environmental changes through time in the seven sectors based on the molluscan records . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 112 eemian species sorted after climatic affinities . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 112 the bælt sea . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 112 the baltic . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 116 the kattegat . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 118 the north sea . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 120 vendsyssel . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 122 skagen . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 125 early/middle weichselian species sorted after climatic affinities . . . . . . . . . . . . . . . . . . . . 126 the kattegat . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 126 vendsyssel . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 127 skagen . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 129 late weichselian species sorted after climatic affinities . . . . . . . . . . . . . . . . . . . . . . . . . 130 vendsyssel . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 130 skagen . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 132 holocene species sorted after climatic affinities . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 133 the bælt sea . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 133 the baltic . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 135 the kattegat . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 138 the limfjord . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 140 the north sea . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 143 vendsyssel . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 146 skagen . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 149 the environmental changes within the seven regions through the late quaternary evaluated by the molluscan communities met with in the seven stages . . . . . . . . . . . . 151 eemian stage 130 000 – 115 000 b.p. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 152 the bælt sea, region 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 152 the baltic, region 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 154 the kattegat, region 3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 154 the north sea, region 5 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 155 the vendsyssel region, region 6 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 156 the skagen region, region 7 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 156 early/middle weichselian stage 115 000 – 25 000 b.p. . . . . . . . . . . . . . . . . . . . . . . . . . 157 the kattegat, region 3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 157 the vendsyssel region, region 6 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 157 geus bulletin no 3.pmd 28-06-2004, 08:454 5 the skagen region, region 7 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 157 late weichselian stage 25 000 – 10 000 b.p. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 158 the vendsyssel region, region 6 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 158 the skagen region, region 7 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 159 the preboreal–boreal stage 10 000 – 8000 14c years b.p. . . . . . . . . . . . . . . . . . . . . . . . 159 the north sea, region 5 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 159 the skagen region, region 7 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 159 the atlantic stage 8000–5000 14c years b.p. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 160 the bælt sea, region 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 160 the baltic, region 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 160 the kattegat, region 3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 160 the limfjord, region 4 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 161 the north sea, region 5 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 161 the vendsyssel region, region 6 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 161 the skagen region, region 7 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 163 the subboreal stage 5000–2500 14c years b.p. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 163 the bælt sea, region 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 163 the baltic, region 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 163 the kattegat, region 3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 164 the limfjord, region 4 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 164 the north sea, region 5 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 166 the vendsyssel region, region 6 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 166 the skagen region, region 7 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 167 the subatlantic stage 2500– 14c years b.p. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 167 the bælt sea, region 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 168 the baltic, region 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 168 the kattegat, region 3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 168 the limfjord, region 4 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 169 the north sea, region 5 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 169 the vendsyssel region, region 6 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 170 the skagen region, region 7 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 171 concluding remarks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 172 acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 172 references . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 173 list of synonyms . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 180 index of species . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 188 appendix 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 197 recent species . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 197 subfossil species . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 201 recent species sorted after climatic affinities . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 204 subfossil species sorted after climatic affinities . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 208 appendix 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 213 appendix 3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 218 appendix 4 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 258 appendix 5 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 260 appendix 6 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 264 geus bulletin no 3.pmd 28-06-2004, 08:455 6 geus bulletin no 3.pmd 28-06-2004, 08:456 7 abstract petersen, k.s. 2004: late quaternary environmental changes recorded in the danish marine molluscan faunas. geological survey of denmark and greenland bulletin 3, 268 pp. late quaternary, marine deposits in denmark have yielded 247 subfossil species of molluscs. the sites are presented, and comparisons are made between the subfossil mollusc assemblages and the 278 shell-bearing mollusc species presently living in the danish seas. 184 species are common to the two groups. the 63 species no longer occurring around denmark are used as indicators of changing environmental conditions, including temperature, salinity and depth, throughout the last 130 000 years. seven modern faunal regional units are defined and considered: the bælt, the baltic, the kattegat, the limfjord, the north sea and the vendsyssel regions, and the skagen area based on the skagen iii well dgu file no. 1.287. the late quaternary, marine, shell-bearing molluscs, comprising 341 subfossil and recent species, are characterised from the point of view of climatic (i.e. arctic, subarctic, boreal and lusitanian) affinities and animal–sediment relationships. on this background the faunal and environmental evolution recorded in the 217 m long skagen well core is analysed and described. the mollusc assemblages in the skagen sequence indicate a deeper-water facies during the eemian, the weichselian and the older holocene in contrast to what hitherto was known in other parts of the danish area during the late quaternary. for the skagen well the chronozones preboreal/boreal, atlantic, subboreal and subatlantic can be identified by 14c dating. the environmental changes within the seven regions through the late quaternary are evaluated by depicting the molluscan communities encountered in the seven late quaternary stages together with remarks on studies of the neighbouring areas. by following the marine communities through the late quaternary in the light of the classical bottom communities sensu c.g.j. petersen, it is demonstrated how facies have changed both through time and space within the danish marine realm. the wellestablished, more temperate eemian marine fauna was closely associated with shallow-water environments. the inferred climatic changes reflect an interglacial–glacial cycle. however, the climatically induced changes during the holocene in the marine environment were small and overshadowed by the facies changes. out of the 341 species recorded in this study, 140 occur in the eemian, 36 in the early/middle weichselian and 41 in the late weichselian. the holocene fauna is represented by 183 species of shell-bearing molluscs, of which the first recorded occurrence of 148 species has been radiocarbon-dated. author’s address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ksp@geus.dk geus bulletin no 3.pmd 28-06-2004, 08:457 8 50 km korsør holbæk høng blåvands huk mandø hølade frederikshavn korupsø agger tange agger bovbjerg ertebølle tastum sø gedser darss stavtrup hals hollerup ejby bro limhamn aarhus stensigmose voderup klint ristinge strandegaard dyrehave hobro aalbæk jydske rev saltholm læsø anholt skagen amager djursland vendsyssel dybvad skærumhede ærø jerup vester holmen samsø højen vust kovad bro vognsbøl fredericia esbjerg varde forballum grærup farup tønder yder bjerrum røjle klint holmstrup røsnæs ulfborg bulbjerg løkken strandby bindslev bornholm møn rügen vejle fjord mariager fjord tybrind vig sidinge fjord skive fjord nissum fjord isefjord roskilde fjord limfjorden fakse bugt skagerrak lille bæ lt store bæ lt ø resund københavn jylland sjælland fyn north sea baltic sea bælt sea kattegat 57° 55° 8° 10° 12° 14° fig. 1. location map with late quaternary marine localities and names of areas on land and of danish waters. geus bulletin no 3.pmd 28-06-2004, 08:458 9 introduction in the middle of 19th century, denmark had its first ‘geology of denmark’ published by g. forchhammer, in 1835. however, as forchhammer expressed it in 1851 when making some notes on the work by the malacologist o.a.l. mørch (1828–1878) at the mineralogical museum of copenhagen. it has hitherto been enough for the geognost to establish formations using the characteristic fossils, but in the future we have to give a closer description from a zoological point of view (petersen 1997, p. 5). considering only the younger deposits, the efforts of the zoologist in geological works are highly significant and became important already in the 19th century. c.g.j. petersen (1860–1928) is an outstanding example of such an influence with his work on the extent of shell-bearing molluscs in the danish seas inside the skagen (petersen 1888, 1893). here he points to the faunal conditions also in the pleistocene and holocene marine deposits compared to the recent distribution. in the description accompanying the geological map sheets of vendsyssel (jessen 1899), jessen gives full credit to c.g.j. petersen and a. jensen (1866– 1953) for their studies on the molluscan species recorded from that part of the country. later both petersen and jensen contributed further to our knowledge of the marine molluscan fauna. petersen formed the concept of the bottom communities (petersen & jensen 1911; petersen 1913, 1914, 1915, 1918) that has been the tool for further work, not only within the danish waters but all over the world with the so-called parallel bottom communities (thorson 1957). though the concept of parallel molluscan communities in the sense of thorson (1957) has been considerably modified in the last 30 years (erwin 1983), there remains a recognition that particular molluscan assemblages are associated with various types of habitat. in 1899 the zoologist v. nordmann (1872–1962) was engaged by the geological survey of denmark to study the molluscs from the quaternary deposits. part of this work was already reflected in the next geological map sheet covering the southern part of vendsyssel (jessen 1905). here nordmann has identified the molluscs and given the faunal remarks on the holocene marine fauna in the north-eastern part of the limfjord (fig. 1). in his work, the zoological considerations are given, elucidating the holocene palaeoenvironments. however, from the beginning of the century nordmann touched upon many other aspects within the late quaternary marine environments which form the most important base for the present study covering marine deposits from the eemian, the weichselian and the holocene. in the following chapter the presentation of some observed sites with marine sediments will be given as an introduction to an answer to the question raised by petersen (1910, p. 29): “what i have often missed in the geological studies is a thorough or detailed comparison between the fossil faunas and the molluscan faunas now living before our eyes”. the aim of this work is to characterise the changing environments in the danish waters through time as seen in the macrofaunas and bottom communities mainly based on molluscs. danish sites with marine sediments initially, the findings and descriptions of the danish marine localities shown in fig. 1 were part of the university studies pioneered by g. forchhammer. however, since the start of the geological survey of denmark in 1888, much of the information has come from the systematic mapping of denmark, and the results have been published in the descriptions to the geological map sheets of denmark (fig. 2). as seen from the plan for the geological mapping of denmark (e.g. sørensen & nielsen 1978) it was decided to do the mapping first in the northern parts of jylland and sjælland and to present a record of the marine deposits from the areas mapped. today, up to 80 per cent of the country has been mapped and descriptions for many map sheets have been published. the main information on the holocene marine molgeus bulletin no 3.pmd 28-06-2004, 08:459 10 luscs is available in these publications and is used in the present description supplemented by specific molluscan studies within the areas. consequently, the frame will be the transition area between the north sea and the baltic and the descrip57° 55° 8° 10° 12° 14° 50 km 1. bælt sea 2. baltic sea 3. kattegat vendsyssel 7. skagen n o r t h s e a 4. limfjorden 5. 6. region 1 – 7 fig. 2. the frame for the seven regions follows mainly the pattern of the old geological map sheets (sørensen & nielsen 1978, fig. 1) and partly the regions used in jensen & knudsen (1995, fig. 1). 1: the bælt sea covering the southern part of the bælts. 2: the baltic covering the southern part of øresund and east of darss–gedser. 3: the kattegat region covering the northern part of the bælts and øresund. 4: the western limfjord – except the north sea coastal region. 5: the north sea with coastal regions and skagerrak. 6: vendsyssel including former marine areas. 7: skagen, mainly the skagen well dgu file no. 1.287. tion mainly based on the geological map sheets found in the following regions shown in fig. 2: (1) the bælt sea; (2) the baltic; (3) the kattegat with bordering fjords; (4) the limfjord; (5) the north sea; (6) vendsyssel and the skagen well iii, dgu file no. 1, 287. the late pleistocene in 1841 forchhammer found the cyprina clay to the southern part of denmark, naming the unit after the dominating bivalve (forchhammer 1842). first, however, forchhammer referred the thick shell molluscs to glossus humanus rather than to arctica islandica. consequently, he placed the deposits in the ‘brunkulsformation’, viz. the tertiary. when finally realising that the common species was arctica islandica, he transferred the deposits to the so-called ‘rullestensformation’, viz. the quaternary. along with the investigations of the cyprina clay through the years since 1841, the actual stratigraphical position was very much under debate, and it was not until 1928 when nordmann wrote his la position stratigraphique des dépôts d’eem that the cyprina clay attained its final position: “appartenant à la dernière période interglaciaire” (nordmann 1928, p. 65). later the name ‘eemian’ became the designation for the whole interglacial, according to gripp (1964, pp. 215–216). johnstrup (1882a) gave the first detailed description of the cyprina clay in denmark and slesvig. also in the northern part of denmark, late pleistocene deposits were studied by johnstrup (1882b), but with references to the earlier works by forchhammer (1822), bredsdorff (1824), faber (1828) and pingel (1828). in 1908 nordmann made his doctoral thesis on the molluscan fauna from the cyprina clay and other central european deposits, forming a part of the publication by madsen et al. (1908). the sequence of interglacial–glacial marine deposgeus bulletin no 3.pmd 28-06-2004, 08:4510 11 its is described from the well at skærumhede (jessen et al. 1910). here the full late pleistocene record is found, although the stratigraphic position was not clear at that time. later investigations, also with studies of the molluscan fauna, were published in 1974 and a late pleistocene age proposed (bahnson et al. 1974). the difference between the boreo-lusitanian community in the boring and the typical eemian community as found in southern denmark was interpreted as difference in facies (bahnson et al. 1974) (see nilsson 1983). in the study of the marine late pleistocene deposits in southern denmark (ødum 1933) based on the record of molluscan species as determined by v. nordmann the finds point to two different deposits in time. one is regarded as eemian and the other as the so-called skærumhede fauna. however, later investigations at strandegaards dyrehave in southern sjælland (petersen & konradi 1974) and at holmstrup in central sjælland (fig. 1; petersen & buch 1974), revealed that the molluscan species found at strandegaards dyrehave, one of the localities of ødum (1933) and regarded as representing the skærumhede fauna, could be eemian but reflecting another facies than the typical eemian on the islands south of fyn. the holmstrup fauna is to be correlated with the arctic marine weichselian in northern jylland which is the upper part of the portlandia arctica zone sensu nordmann (madsen et al. 1908) or the macoma calcarea zone sensu petersen (bahnson et al. 1974, fig. 7), see fig. 3 for stratigraphical position. the aminostratigraphic investigations of the danish late pleistocene deposits as published by miller & mangerud (1985) sustain only to some extent the abovementioned correlations: “none of the sites regarded here as eemian (strandegaards dyrehave) gave ratios as high as in holsteinian deposits or as low as in middle weichselian deposits” (miller & mangerud 1985, p. 261). in the case of the holmstrup weichselian site, only three out of eleven individuals of macoma calcarea gave weichselian ratios (miller & mangerud 1985, p. 264). the marine molluscan fauna of the late weichselian has been studied intensively only from the vendh ol oc en e la te w ei ch se lia n pl ei st oc en e m id dl e w ei ch se lia n ea rl y w ei ch . eemian l a t e q u a t e r n a r y subatlantic subboreal atlantic boreal preboreal younger dryas allerød older dryas bølling historical age iron age bronze age neolithic mesolithic palaeolithic young baltic swedish old baltic norwegian 1000 2000 3000 4000 5000 6000 7000 8000 9000 10000 11000 12000 13000 14000 15000 16000 17000 18000 19000 20000 21000 22000 1050 2050 2900 3700 4400 5500 6100 7200 8100 9100 10000 10200 11100 12100 12800 13400 14200 15100 16100 17000 17600 18500 25000 75000 115000 130000 period epoch age culture stagechron ice stream chronology calendar years bp 14c years bp fig. 3. stratigraphic framework for the late quaternary deposits from pedersen & petersen (1997). geus bulletin no 3.pmd 28-06-2004, 08:4511 12 syssel area recorded in the publications by jessen (1899, 1936). nearly 30 molluscan samples from these late weichselian – younger yoldia sea deposits have been dated (krog & tauber 1974). theevaluationof themolluscancommunities invendsyssel reveals the changing late weichselian sea level (petersen 1984), and the highest marine shoreline, around 60 m a.s.l., in northern denmark can be shown to develop between 14 000 and 13 000 b.p. (14c years). the holocene forchhammer participated in the work of the so-called ‘lejrekomité’, an interdisciplinary committee studying human remains along the shore. this commission gave the first – and now famous – description of the ‘køkkenmødding’ (kitchen midden), a mound consisting of shells of edible molluscs and other refuse, marking the site of a prehistoric human habitation (hanks 1971). ‘køkkenmødding’ is one of the few danish international terms (forchhammer et al. 1851). the work of the ‘lejrekomité’ was concentrated on the marine molluscs in order to establish out whether the shell deposits were naturally based – oyster banks – or whether they were formed as waste deposits produced by men living at coastal sites. the other members of the commission were j. worsaae and j. steenstrup, representing archaeology and zoology respectively. consistently, the study of the molluscan elements was based mainly on steenstrup’s work. however, while working in the commission, forchhammer continued his studies on the sea levels (forchhammer 1838, 1840). this was essential for the discussion of whether the molluscs found belonged to raised marine deposits or were gathered by man. forchhammer’s study led to the concept of raised marine deposits north of a line from nissum fjord to south of korsør in the storebælt area (fig. 1). this line still carries the name of forchhammer and divides the country into two parts, with the raised marine areas to the north-east, and to the south-west the area where the land has been sinking. together with the study of the holocene molluscan fauna by johnstrup (1882b), such observations on shorelines were also collected. it became one of the points specially mentioned in the instructions for the autographic geologists when the systematic geological mapping of denmark was started in 1888 by the geological survey of denmark (sørensen & nielsen 1978). the recent fauna of shell-bearing molluscs compared to the subfossil fauna the record of recent danish shell-bearing molluscs has been taken from the annotated check list of recent marine molluscs of danish waters (jensen & knudsen 1995). in appendix 1 the species are presented taxonomically following jensen & knudsen (1995). late immigrants from the last centuries – transferred by man -– have been omitted from the list, because the aim of the present study is to present the development in the subfossil late quaternary molluscan fauna also in appendix 1 compared to the natural fauna of today. according to fredén (1986), subfossil means that the weight of the object when found does not exceed its original weight, which is obviously the case for younger deposits seen geologically as shells from the late quaternary. in all, 278 recent species of shell-bearing molluscs are recorded from the danish waters: the class polyplacophora is represented by seven geus bulletin no 3.pmd 28-06-2004, 08:4512 13 species forming 2.5% of the total number of known species. the class gastropoda is represented by 151 species forming 54.3% of the total number of known species. the class scaphopoda is represented by three species forming 1.1% of the total number of known species. the class bivalvia is represented by 117 species forming 42.1% of the total number of known species. the list of known finds of subfossil species amounts to 247 species. with regard to the classes, it appears that polyplacophora is now represented by only one species, which formed 0.4% of the total subfossil molluscan species. within the class gastropoda 125 species occur, forming 50.6% of the total number of subfossil species, a figure which is nearly 5% lower than that for recent gastropods. the class scaphopoda is represented by five fossil species which form 2.0% of the subfossil shell-bearing species which is a little higher than the ratio for the recent fauna. the class bivalvia is represented by 116 species forming 47.0% of the total, which is a little more than 7% above the recent ratio. the low number of subfossil polyplacophora can be explained by the fact that the shells from those species are nearly always broken, and this excludes identification to species level, so to say, following the statement made by knudsen (1970, p. 1): “isolated and worn plates were neglected altogether”. among the gastropods, the subclasses and orders, except the order heterostropha within the subclass heterobranchia, have a lower representation of subfossil finds than of recent ones. the heterostropha, which has a 2.5% higher representation among the subfossil finds than among the recent ones, is a group fig. 4. regional division of the european seas from feyling-hanssen (1955). geus bulletin no 3.pmd 28-06-2004, 08:4513 14 of mostly tiny specimens which might be more looked for in the geological samples than in the recent bottom samples often used in the more practical work of evaluation benthos introduced by c.g.j.petersen. however, many of these small species should be considered with the utmost care, with respect to the difficulty of identifying them to species level within subfossil material. the reason why the class scaphopoda has a twice as great a representation within the subfossil material cannot be given, although it is tempting to regard the different palaeoenvironment back in the late quaternary as the explanation of the higher frequency. the greater variety of palaeoenvironment and different climate back in time is clearly the reason why the bivalvia within all subclasses has a higher percentage than in the recent fauna. however, as an overview, the total subfossil species could be compared to the recent ones arranged also after their climatic affinities, as will be thoroughly discussed in one of the following chapters. with respect to distribution of molluscan species within the north atlantic – west european realm, four zones may be distinguished, viz.: the arctic = a, the subarctic = s, the boreal = b and the lusitanian = l (figs 4, 5). it appears from the comparison between subfossil species and recent species sorted after climatic affinity (appendix 1) that the subfossil species have their dominance in the extreme groups, i.e. arctic = a; arctic/ subarctic = as; arctic, subarctic and boreal = asb and subarctic/boreal, while the species with a wide tolerance – arctic, subarctic, boreal and lusitanian = asbl – have a higher representation within the recent fauna. also the middle group, which is represented by faunal element from the subarctic, boreal and lusitanian, the boreal and lusitanian (which is the most numerous group with 140 subfossil species) has a higher representation in the recent fauna. but the group of purely lusitanian species has a clearly better representation among the subfossil species, as seen by the percentage figure 6.2% compared to 0.7% for the purely lusitanian faunal elements among the subfossil and recent faunas respectively. these observations reveal that the late quaternary fauna covers a period of 130 000 years with changing climatic conditions both with colder and warmer periods than at present. so considering the totals of subfossil and recent species one has to discuss the difference not only quantitatively but qualitatively; because only 184 species are shared between the late quaternary and the recent finds, while 63 species have to be considered as particular ones occurring within the late quaternary during the eemian, the weichselian or the holocene, in one, two or in all three groups but not the recent one. within the bivalvia, the highest amount of subfossil species (31) found only in the late quaternary occur. such species are also the species which must be focused on in the evaluation of the changing environment through time. fig. 5. regional division of the european seas according to símonarson et al. (1998). geus bulletin no 3.pmd 28-06-2004, 08:4514 15 molluscan finds within the seven regions during the holocene the molluscan finds within each region (see fig. 2) from the holocene, as appearing mainly from the descriptions accompanying the geological map sheets of denmark, are presented. the bælt sea area from the bælt sea area the information on the occurrences of molluscs has been taken from the following map sheets: madsen (1902) and jessen (1907 (contributions by v.nordmann), 1935, 1945); v.milthers (1940) and k. milthers (1959). nordmann (1906) has a record of molluscs found in skælskør nor (sw sjælland) and petersen records from the areas south of fyn, storebælt and lillebælt (1985c, 1989). subfossil holocene species in the bælt sea area class gastropoda subclass prosobranchia order neotaenioglossa littorina littorea (linnaeus 1758) littorina obtusata (linnaeus 1758) littorina saxatilis (olivi 1792) littorina tenebrosa (montagu 1803) lacuna pallidula (da costa 1778) lacuna vincta (montagu 1803) hydrobia ulvae (pennant 1777) hydrobia ventrosa (montagu 1803) onoba semicostata (montagu 1803) rissoa albella lovén 1846 rissoa inconspicua alder 1844 rissoa membranacea (j. adams 1800) bittium reticulatum (da costa 1778) order heterogastropoda triphora adversa (montagu 1803) order neogastropoda hinia reticulata (linnaeus 1758) subclass heterobranchia order heterostropha omalogyra atomus (phillippi 1841) odostomia conoidea winckworth 1932 subclass opisthobranchia order anaspidea retusa obtusa (montagu 1803) retusa truncatula (bruguière 1792) akera bullata müller 1776 subclass pulmonata order basommatophora lymnaea peregra (müller 1774) class bivalvia subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 modiolula phaseolina (philippi 1844) modiolus modiolus (linnaeus 1758) musculus discors (linnaeus 1767) order pterioida ostrea edulis linnaeus 1758 subclass heterodonta order veneroida mysella bidentata (montagu 1803) tridonta borealis schumacher 1817 parvicardium exiguum (gmelin 1791) parvicardium ovale (sowerby 1840) parvicardium scabrum (philippi 1844) cerastoderma edule (linnaeus 1758) spisula subtruncata (da costa 1778) angulus tenuis (da costa 1778) macoma balthica (linnaeus 1758) scrobicularia plana (da costa 1778) abra alba (wood 1802) arctica islandica (linnaeus 1767) paphia aurea (gmelin 1791) tapes decussatus (linnaeus 1758) venerupis pullastra (montagu 1803) order myoida mya arenaria linnaeus 1758 mya truncata linnaeus 1758 corbula gibba (olivi 1792) hiatella arctica (linnaeus 1758) barnea candida (linnaeus 1758) zirfaea crispata (linnaeus 1758) total for the holocene bælt sea: 47 (19.0%) geus bulletin no 3.pmd 28-06-2004, 08:4515 16 the baltic area the baltic area is here restricted to the area east of darss and south of øresund at saltholm, which must be considered the baltic sensu stricto when regarding the present distribution of the marine fauna and also taking into consideration the subfossil holocene molluscan fauna, as will be demonstrated by a following comparison with the other areas. the main map sheet published is by v. milthers from 1908 with contributions by v. nordmann on the holocene molluscan fauna. the subfossil holocene fauna has also been studied later in the western part by petersen (1994b). in the description accompanying the map sheet bornholm (grönwall & milthers 1916) there is no record of a mollusc fauna. subfossil holocene species in the baltic area class gastropoda subclass prosobranchia order neotaenioglossa littorina littorea (linnaeus 1758) littorina tenebrosa (montagu 1803) hydrobia ulvae (pennant 1777) hydrobia ventrosa (montagu 1803) rissoa albella lovén 1846 rissoa inconspicua alder 1844 rissoa membranacea (j. adams 1800) bittium reticulatum (da costa 1778) aporrhais pespelicani (linnaeus 1758) order neogastropoda hinia reticulata (linnaeus 1758) subclass opisthobranchia order anaspidea retusa truncatula (bruguière 1792) subclass pulmonata order basommatophora lymnaea peregra (müller 1774) class bivalvia subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 subclass heterodonta order veneroida parvicardium exiguum (gmelin 1791) cerastoderma edule (linnaeus 1758) cerastoderma glaucum (poiret 1789) macoma balthica (linnaeus 1758) scrobicularia plana (da costa 1778) order myoida corbula gibba (olivi 1792) total for the holocene baltic: 19 (7.7%) the kattegat area with fjords the kattegat area sensu lato includes the fjords, i.e. the northern part of the lillebælt area, storebælt and øresund. therefore the following map sheets are taken within this area: 1. the north-eastern part of sjælland described by rørdam (1893), who published a detailed description of the holocene marine deposits from northeast sjælland already in 1891 and continued with the description of the map sheet københavn and roskilde (1899) where the southernmost parts of the roskilde fjord and the øresund are described in great detail for the marine holocene part. 2. furthermore, rørdam & v. milthers published the description for the geological map sheet of nw sjælland in 1900 and nordmann on the molluscs in sidinge fjord (westerby 1933). 3. from the north-western part of fyn and the island of samsø by madsen (1897, 1900) and together with ussing for the north-eastern part of fyn (ussing & madsen 1897). 4. the map sheet of fredericia (nordmann 1958) covers the northern part of the lillebælt and vejle fjord on the eastern coast of jylland. 5. from the islands of læsø and anholt in the kattegat the description was given by jessen (1897) and nordmann (1903a). 6. from hobro by nordmann (jessen 1927). 7. from mariager fjord by nordmann (ødum 1929). 8. the peninsula of djursland has been mapped during the last decades, and the description of the holocene marine molluscan fauna is by petersen (pedersen & petersen 1997) and petersen (1993). subfossil holocene species in the kattegat class gastropoda subclass prosobranchia order archaeogastropoda theodoxus fluviatilis (linnaeus 1758) order neotaenioglossa littorina littorea (linnaeus 1758) littorina obtusata (linnaeus 1758) littorina saxatilis (olivi 1792) littorina tenebrosa (montagu 1803) lacuna vincta (montagu 1803) hydrobia ulvae (pennant 1777) hydrobia ventrosa (montagu 1803) skeneopsis planorbis (fabricius 1780) onoba semicostata (montagu 1803) onoba vitrea (montagu 1803) geus bulletin no 3.pmd 28-06-2004, 08:4516 17 rissoa albella lovén 1846 rissoa inconspicua alder 1844 rissoa membranacea (j. adams 1800) bittium reticulatum (da costa 1778) order heterogastropoda triphora adversa (montagu 1803) order neogastropoda buccinum undatum linnaeus 1758 hinia reticulata (linnaeus 1758) subclass heterobranchia order heterostropha omalogyra atomus (phillippi 1841) brachystomia eulimoides hanley 1844 chrysallida spiralis (montagu 1803) subclass opisthobranchia order anaspidea retusa obtusa (montagu 1803) retusa truncatula (bruguière 1792) akera bullata müller 1776 class bivalvia subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 musculus discors (linnaeus 1767) order pterioida heteranomia squamula (linnaeus 1758) ostrea edulis linnaeus 1758 subclass heterodonta order veneroida mysella bidentata (montagu 1803) acanthocardia echinata (linnaeus 1758) parvicardium exiguum (gmelin 1791) parvicardium scabrum (philippi 1844) cerastoderma edule (linnaeus 1758) cerastoderma glaucum (poiret 1798) macoma balthica (linnaeus 1758) scrobicularia plana (da costa 1778) abra alba (wood 1802) arctica islandica (linnaeus 1767) paphia aurea (gmelin 1791) tapes decussatus (linnaeus 1758) venerupis pullastra (montagu 1803) order myoida mya arenaria linnaeus 1758 mya truncata linnaeus 1758 corbula gibba (olivi 1792) hiatella rugosa (linnaeus 1758) total for the holocene kattegat: 45 (18.2%) the limfjord area from the limfjord area (western part), excluding the part which falls within vendsyssel, only one description for a map sheet has been published (gry 1979). however, the molluscs are recorded in publications by petersen (1976, 1981, 1985a, 1986a) and in rasmussen & petersen (1980). furthermore, v.nordmann collected holocene marine shells from the western limfjord in 1902–1903 which were further examined by erna nordmann and leifur símonarson in the sixties as mentioned in petersen (1976, p. 78). it must be emphasised that c.g.j. petersen in 1888 discussed the subfossil fauna also from the limfjord, which was earlier the topic of collin (1884). subfossil holocene species in the limfjord class gastropoda subclass prosobranchia order archaeogastropoda patella vulgata linnaeus 1758 helcion pellucidum (linnaeus 1758) iothia fulva (müller 1776) acmaea tessulata (müller 1776) acmaea virginea (müller 1776) margarites helicinus (phipps 1774) gibbula cineraria (linnaeus 1758) gibbula tumida (montagu 1803) skenea serpuloides (montagu 1808) skenea basistriata (jeffreys 1877) order neotaenioglossa littorina littorea (linnaeus 1758) littorina obtusata (linnaeus 1758) littorina saxatilis (olivi 1792) littorina tenebrosa (montagu 1803) lacuna pallidula (da costa 1778) lacuna parva (montagu 1803) lacuna vincta (montagu 1803) hydrobia ulvae (pennant 1777) hydrobia ventrosa (montagu 1803) skeneopsis planorbis (fabricius 1780) alvania lactea (michaud 1830) alvania punctura (montagu 1803) cingula semistriata (montagu 1808) onoba semicostata (montagu 1803) onoba proxima (forbes & hanley 1850) onoba vitrea (montagu 1803) rissoa albella lovén 1846 rissoa inconspicua alder 1844 rissoa membranacea (j. adams 1800) rissoa parva (da costa 1779) rissoa violacea desmarest 1814 geus bulletin no 3.pmd 28-06-2004, 08:4517 18 caecum glabrum (montagu 1803) bittium reticulatum (da costa 1778) turritella communis risso 1826 aporrhais pespelicani (linnaeus 1758) lunatia alderi (forbes 1838) lunatia catena (da costa 1778) order heterogastropoda triphora adversa (montagu 1803) cerithiopsis barleei (jeffreys 1867) cerithiopsis tubercularis (montagu 1803) epitonium clathrus (linnaeus 1758) epitonium turtonis (turton 1819) aclis minor (brown 1827) vitreolina philippii (rayneval & ponzi 1854) order neogastropoda nucella lapillus (linnaeus 1758) buccinum undatum linnaeus 1758 hinia incrassata (ström 1768) hinia pygmaea (lamarck 1822) hinia reticulata (linnaeus 1758) oenopota turricola (montagu 1803) raphitoma purpurea (montagu 1803) raphitoma linearis (montagu 1803) subclass heterobranchia order heterostropha omalogyra atomus (phillippi 1841) brachystomia eulimoides hanley 1844 odostomia scalaris macgillivray 1843 chrysallida decussata (montagu 1803) chrysallida eximia (jeffreys 1849) chrysallida indistincta (montagu 1808) chrysallida obtusa (brown 1827) chrysallida spiralis (montagu 1803) ebala nitidissima (montagu 1803) eulimella laevis (brown 1827) eulimella scillae (scacchi 1835) ondina divisa (j. adams 1797) ondina diaphana (jeffreys 1848) odostomia acuta jeffreys 1848 odostomia conoidea winckworth 1932 odostomia turrita hanley 1844 odostomia albella lovén 1846 odostomia plicata (montagu 1803) turbonilla crenata (brown 1827) turbonilla delicata (monterosato 1874) turbonilla lactea (linnaeus 1758) subclass opisthobranchia order bullomorpha acteon tornatilis (linnaeus 1758) cylichna cylindracea (pennant 1777) cylichna alba (brown 1827) philine aperta (linnaeus 1767) philine punctata (adams 1800) order anaspidea diaphana minuta brown 1827 retusa obtusa (montagu 1803) retusa truncatula (bruguière 1792) retusa umbilicata (montagu 1803) akera bullata müller 1776 class bivalvia subclass palaeotaxodonta order nuculoida nucula nitidosa winckworth 1930 nucula nucleus (linnaeus 1767) nuculoma tenuis (montagu 1808) subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 modiolula phaseolina (philippi 1844) modiolus adriaticus (lamarck 1819) modiolus modiolus (linnaeus 1758) musculus discors (linnaeus 1767) modiolaria tumida (hanley 1843) order pterioida aequipecten opercularis (linnaeus 1758) chlamys varia (linnaeus 1758) delectopecten vitreus (gmelin 1791) palliolum striatum (müller 1776) palliolum tigerinum (müller 1776) pododesmus patelliformis (linnaeus 1761) anomia ephippium linnaeus 1758 heteranomia squamula (linnaeus 1758) ostrea edulis linnaeus 1758 subclass heterodonta order veneroida lucinoma borealis (linnaeus 1758) thyasira flexuosa (montagu 1803) mysella bidentata (montagu 1803) tellimya ferruginosa (montagu 1803) turtonia minuta (fabricius 1780) lepton nitidum (turton 1822) tridonta borealis schumacher 1817 acanthocardia echinata (linnaeus 1758) parvicardium exiguum (gmelin 1791) parvicardium ovale (sowerby 1840) parvicardium scabrum (philippi 1844) cerastoderma edule (linnaeus 1758) cerastoderma glaucum (poiret 1789) mactra stultorum (linnaeus 1758) lutraria lutraria (linnaeus 1758) spisula elliptica (brown 1827) spisula solida (linnaeus 1758) geus bulletin no 3.pmd 28-06-2004, 08:4518 19 spisula subtruncata (da costa 1778) ensis ensis (linnaeus 1758) phaxas pellucidus (pennant 1777) angulus tenuis (da costa 1778) fabulina fabula (gmelin 1791) macoma balthica (linnaeus 1758) donax vittatus (da costa 1778) gari fervensis (gmelin 1791) scrobicularia plana (da costa 1778) abra alba (wood 1802) abra nitida (müller 1776) abra prismatica (montagu 1803) arctica islandica (linnaeus 1767) chamelea striatula (da costa 1778) clausinella fasciata (da costa 1778) paphia aurea (gmelin 1791) tapes decussatus (linnaeus 1758) timoclea ovata (pennant 1777) venerupis rhomboides (pennant 1777) venerupis pullastra (montagu 1803) mysia undata (pennant 1777) order myoida mya truncata linnaeus 1758 corbula gibba (olivi 1792) hiatella arctica (linnaeus 1758) saxicavella jeffreysi winckworth 1930 barnea candida (linnaeus 1758) pholas dactylus linnaeus 1758 zirfaea crispata (linnaeus 1758) subclass anomalodesmata order pholadomyoida thracia phaseolina (lamarck 1818) total for the holocene limfjord: 147 (59.5%) the north sea in the north sea region the map sheet blaavands huk (fig. 1) forms the southernmost part of what is covered by the present presentation regarding the holocene deposits, and this area was described by jessen (1925). nordmann (in jessen 1925) contributed with the study of the molluscs. 1. in the work by petersen (1985a) the molluscan fauna in the coastal region – the aggertange – is recorded. 2. the geological map sheet from ulfborg was published by petersen et al. (1992a), and the molluscan fauna treated by petersen, but not yet published, is included. 3. in 1994 the holocene molluscs from the jydske rev were studied and reported in a work for the danish coastal authority (petersen 1994a), and with minor corrections published in petersen (1998). subfossil holocene species in the north sea class gastropoda subclass prosobranchia order archaeogastropoda gibbula cineraria (linnaeus 1758) theodoxus fluviatilis (linnaeus 1758) order neotaenioglossa littorina littorea (linnaeus 1758) littorina obtusata (linnaeus 1758) littorina saxatilis (olivi 1792) lacuna pallidula (da costa 1778) lacuna parva (montagu 1803) lacuna vincta (montagu 1803) hydrobia ulvae (pennant 1777) hydrobia ventrosa (montagu 1803) cingula turgida (jeffreys 1870) onoba vitrea (montagu 1803) rissoa albella lovén 1846 rissoa inconspicua alder 1844 rissoa membranacea (j. adams 1800) rissoa violacea desmarest 1814 caecum glabrum (montagu 1803) bittium reticulatum (da costa 1778) turritella communis risso 1826 aporrhais pespelicani (linnaeus 1758) lunatia alderi (forbes 1838) lunatia catena (da costa 1778) order heterogastropoda triphora adversa (montagu 1803) epitonium clathrus (linnaeus 1758) aclis ascaris (turton 1819) aclis minor (brown 1827) aclis walleri jeffreys 1867 order neogastropoda nucella lapillus (linnaeus 1758) buccinum undatum linnaeus 1758 hinia pygmaea (lamarck 1822) hinia reticulata (linnaeus 1758) subclass heterobranchia order heterostropha brachystomia eulimoides hanley 1844 chrysallida indistincta (montagu 1808) chrysallida spiralis (montagu 1803) eulimella laevis (brown 1827) ondina diaphana (jeffreys 1848) odostomia conoidea winckworth 1932 geus bulletin no 3.pmd 28-06-2004, 08:4519 20 odostomia albella lovén 1846 turbonilla crenata (brown 1827) turbonilla delicata (monterosato 1874) turbonilla lactea (linnaeus 1758) subclass opisthobranchia order bullomorpha acteon tornatilis (linnaeus 1758) cylichna alba (brown 1827) order anaspidea retusa obtusa (montagu 1803) retusa truncatula (bruguière 1792) retusa umbilicata (montagu 1803) class bivalvia subclass palaeotaxodonta order nuculoida nucula nitidosa winckworth 1930 nucula nucleus (linnaeus 1767) nucula sulcata (bronn 1831) nuculoma tenuis (montagu 1808) subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 musculus discors (linnaeus 1767) order pterioida chlamys varia (linnaeus 1758) heteranomia squamula (linnaeus 1758) ostrea edulis linnaeus 1758 subclass heterodonta order veneroida thyasira flexuosa (montagu 1803) mysella bidentata (montagu 1803) tellimya ferruginosa (montagu 1803) lepton nitidum (turton 1822) acanthocardia echinata (linnaeus 1758) parvicardium exiguum (gmelin 1791) parvicardium ovale (sowerby 1840) parvicardium scabrum (philippi 1844) parvicardium minimum (philippi 1836) cerastoderma edule (linnaeus 1758) mactra stultorum (linnaeus 1758) spisula elliptica (brown 1827) spisula solida (linnaeus 1758) spisula subtruncata (da costa 1778) ensis ensis (linnaeus 1758) phaxas pellucidus (pennant 1777) angulus tenuis (da costa 1778) fabulina fabula (gmelin 1791) macoma balthica (linnaeus 1758) donax vittatus (da costa 1778) scrobicularia plana (da costa 1778) abra alba (wood 1802) abra nitida (müller 1776) arctica islandica (linnaeus 1767) chamelea striatula (da costa 1778) clausinella fasciata (da costa 1778) paphia aurea (gmelin 1791) tapes decussatus (linnaeus 1758) timoclea ovata (pennant 1777) venerupis pullastra (montagu 1803) dosinia exoleta (linnaeus 1758) dosinia lincta (montagu 1803) order myoida mya truncata linnaeus 1758 corbula gibba (olivi 1792) hiatella arctica (linnaeus 1758) saxicavella jeffreysi winckworth 1930 barnea candida (linnaeus 1758) zirfaea crispata (linnaeus 1758) subclass anomalodesmata order pholadomyoida cochlodesma praetenue (pulteney 1799) thracia phaseolina (lamarck 1818) total for the holocene north sea: 95 (38.5%) the vendsyssel area the vendsyssel area includes the description accompanying the map sheets over the northern, central and southern parts, all by jessen (1899, 1905), but with a contribution by v. nordmann, who wrote the part on the holocene molluscan fauna in the latter publication. in this description by nordmann he presents the different faunal communities as discovered in the subfossil assemblages. it was nordmann’s intention to continue the work further west into the western limfjord area, but his first investigations were not used in the systematic geological mapping. they were, however, of great importance for the understanding of the development of the holocene molluscan fauna (nordmann 1910, 1918). in 1928 in connection with the international congress in copenhagen a final overview by nordmann of the quaternary marine deposits in denmark was given in the summary of the geology of denmark (madsen et al. 1928). here nordmann points to the dosinia layers first described at the beginning of the century from vendsyssel (nordmann 1904), with a record of a fauna not found in the older tapes beds originally demonstrated by petersen (1888). geus bulletin no 3.pmd 28-06-2004, 08:4520 21 later investigations by lauersen (1937) and petersen (1990,1991a, b, 1992) on the dosinia beds at strandby are also included in the list of holocene marine molluscs from vendsyssel. in just pedersen’s thesis on holocene molluscs from 1976 (unpublished), a new fauna element donax vittatus in the dosinia beds is recorded from frederikshavn. subfossil holocene species in the vendsyssel class gastropoda subclass prosobranchia order archaeogastropoda helcion pellucidum (linnaeus 1758) acmaea virginea (müller 1776) gibbula cineraria (linnaeus 1758) gibbula tumida (montagu 1803) order neotaenioglossa littorina littorea (linnaeus 1758) littorina obtusata (linnaeus 1758) littorina saxatilis (olivi 1792) littorina tenebrosa (montagu 1803) lacuna pallidula (da costa 1778) lacuna parva (montagu 1803) lacuna vincta (montagu 1803) hydrobia ulvae (pennant 1777) skeneopsis planorbis (fabricius 1780) alvania lactea (michaud 1830) alvania cimicoides (forbes 1844) alvania punctura (montagu 1803) cingula semistriata (montagu 1808) onoba semicostata (montagu 1803) onoba vitrea (montagu 1803) rissoa albella lovén 1846 rissoa inconspicua alder 1844 rissoa membranacea (j. adams 1800) rissoa parva (da costa 1779) rissoa violacea desmarest 1814 caecum glabrum (montagu 1803) bittium reticulatum (da costa 1778) turritella communis risso 1826 aporrhais pespelicani (linnaeus 1758) trivia monacha (da costa 1778) lunatia alderi (forbes 1838) lunatia catena (da costa 1778) order heterogastropoda triphora adversa (montagu 1803) epitonium clathrus (linnaeus 1758) epitonium turtonis (turton 1819) vitreolina philippii (rayneval & ponzi 1854) order neogastropoda nucella lapillus (linnaeus 1758) buccinum undatum linnaeus 1758 neptunea antiqua (linnaeus 1758) hinia incrassata (ström 1768) hinia pygmaea (lamarck 1822) hinia reticulata (linnaeus 1758) cytharella coarctata (forbes 1840) oenopota turricola (montagu 1803) raphitoma linearis (montagu 1803) subclass heterobranchia order heterostropha omalogyra atomus (phillippi 1841) brachystomia eulimoides hanley 1844 odostomia scalaris macgillivray 1843 chrysallida indistincta (montagu 1808) chrysallida obtusa (brown 1827) chrysallida spiralis (montagu 1803) ebala nitidissima (montagu 1803) eulimella laevis (brown 1827) odostomia conoidea winckworth 1932 odostomia turrita hanley 1844 odostomia albella lovén 1846 odostomia plicata (montagu 1803) turbonilla lactea (linnaeus 1758) subclass opisthobranchia order bullomorpha acteon tornatilis (linnaeus 1758) cylichna cylindracea (pennant 1777) philine aperta (linnaeus 1767) philine punctata (adams 1800) order anaspidea diaphana minuta brown 1827 retusa obtusa (montagu 1803) retusa truncatula (bruguière 1792) retusa umbilicata (montagu 1803) akera bullata müller 1776 class bivalvia subclass palaeotaxodonta order nuculoida nucula nitidosa winckworth 1930 nucula nucleus (linnaeus 1767) subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 modiolula phaseolina (philippi 1844) modiolus adriaticus (lamarck 1819) modiolus modiolus (linnaeus 1758) musculus discors (linnaeus 1767) modiolaria tumida (hanley 1843) order pterioida aequipecten opercularis (linnaeus 1758) geus bulletin no 3.pmd 28-06-2004, 08:4521 22 chlamys varia (linnaeus 1758) pecten maximus (linnaeus 1758) pododesmus patelliformis (linnaeus 1761) anomia ephippium linnaeus 1758 heteranomia squamula (linnaeus 1758) ostrea edulis linnaeus 1758 subclass heterodonta order veneroida lucinoma borealis (linnaeus 1758) thyasira flexuosa (montagu 1803) mysella bidentata (montagu 1803) tellimya ferruginosa (montagu 1803) turtonia minuta (fabricius 1780) lepton nitidum (turton 1822) kellia suborbicularis (montagu 1803) acanthocardia echinata (linnaeus 1758) parvicardium exiguum (gmelin 1791) parvicardium ovale (sowerby 1840) parvicardium scabrum (philippi 1844) cerastoderma edule (linnaeus 1758) laevicardium crassum (gmelin 1791) mactra stultorum (linnaeus 1758) lutraria lutraria (linnaeus 1758) spisula elliptica (brown 1827) spisula solida (linnaeus 1758) spisula subtruncata (da costa 1778) ensis ensis (linnaeus 1758) phaxas pellucidus (pennant 1777) angulus tenuis (da costa 1778) fabulina fabula (gmelin 1791) macoma balthica (linnaeus 1758) macoma calcarea (gmelin 1791) donax vittatus (da costa 1778) gari depressa (pennant 1777) gari fervensis (gmelin 1791) scrobicularia plana (da costa 1778) abra alba (wood 1802) abra nitida (müller 1776) abra prismatica (montagu 1803) arctica islandica (linnaeus 1767) chamelea striatula (da costa 1778) clausinella fasciata (da costa 1778) paphia aurea (gmelin 1791) tapes decussatus (linnaeus 1758) timoclea ovata (pennant 1777) venerupis rhomboides (pennant 1777) venerupis pullastra (montagu 1803) dosinia exoleta (linnaeus 1758) dosinia lincta (montagu 1803) mysia undata (pennant 1777) order myoida mya arenaria linnaeus 1758 mya truncata linnaeus 1758 corbula gibba (olivi 1792) hiatella arctica (linnaeus 1758) hiatella rugosa (linnaeus 1758) saxicavella jeffreysi winckworth 1930 barnea candida (linnaeus 1758) pholas dactylus linnaeus 1758 zirfaea crispata (linnaeus 1758) subclass anomalodesmata order pholadomyoida thracia phaseolina (lamarck 1818) total for holocene vendsyssel: 133 (53.8%) the skagen well area the hitherto recorded molluscan assemblages from danish deposits of late quaternary age are littoral to sublittoral – mostly – especially from the holocene. the new information from the skagen well containing deeper-water deposits is presented below. subfossil holocene species in the skagen well class gastropoda subclass prosobranchia order neotaenioglossa lacuna pallidula (da costa 1778) hydrobia ulvae (pennant 1777) barleeia unifasciata (montagu 1803) onoba vitrea (montagu 1803) rissoa albella lovén 1846 rissoa violacea desmarest 1814 bittium reticulatum (da costa 1778) turritella communis risso 1826 aporrhais pespelicani (linnaeus 1758) lunatia alderi (forbes 1838) lunatia montagui (forbes 1838) order heterogastropoda epitonium trevelyanum (johnston 1841) aclis minor (brown 1827) polygireulima sinuosa (sacco 1836) vitreolina collensi (sykes 1903) vitreolina philippii (rayneval & ponzi 1854) graphis albida (kanmacher 1798) melanella lubrica (monterosato 1891) melanella alba (da costa 1778) hemiaclis ventrosa (jeffreys ms fricle 1874) order neogastropoda buccinum undatum linnaeus 1758 geus bulletin no 3.pmd 28-06-2004, 08:4522 23 hinia pygmaea (lamarck 1822) hinia reticulata (linnaeus 1758) oenopota turricola (montagu 1803) mangelia brachystoma (philippi 1844) subclass heterobranchia order heterostropha chrysallida decussata (montagu 1803) eulimella scillae (scacchi 1835) odostomia conoidea winckworth 1932 odostomia umbilicaris (malm 1863) turbonilla delicata (monterosato 1874) turbonilla sinuosa (jeffreys 1884) subclass opisthobranchia order bullomorpha cylichna alba (brown 1827) order anaspidea retusa truncatula (bruguière 1792) retusa umbilicata (montagu 1803) class bivalvia subclass palaeotaxodonta order nuculoida nucula nitidosa winckworth 1930 nucula nucleus (linnaeus 1767) nuculana minuta (müller 1776) subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 musculus discors (linnaeus 1767) order pterioida chlamys varia (linnaeus 1758) heteranomia squamula (linnaeus 1758) ostrea edulis linnaeus 1758 subclass heterodonta order veneroida thyasira flexuosa (montagu 1803) mysella bidentata (montagu 1803) tellimya ferruginosa (montagu 1803) turtonia minuta (fabricius 1780) acanthocardia echinata (linnaeus 1758) parvicardium minimum (philippi 1836) mactra stultorum (linnaeus 1758) spisula subtruncata (da costa 1778) phaxas pellucidus (pennant 1777) angulus tenuis (da costa 1778) fabulina fabula (gmelin 1791) tellina pygmaea (lovén 1846) donax vittatus (da costa 1778) gari fervensis (gmelin 1791) abra alba (wood 1802) abra nitida (müller 1776) abra prismatica (montagu 1803) arctica islandica (linnaeus 1767) chamelea striatula (da costa 1778) timoclea ovata (pennant 1777) order myoida mya arenaria linnaeus 1758 corbula gibba (olivi 1792) hiatella arctica (linnaeus 1758) saxicavella jeffreysi winckworth 1930 barnea candida (linnaeus 1758) pholas dactylus linnaeus 1758 subclass anomalodesmata order pholadomyoida lyonsia norvegica (gmelin 1791) cochlodesma praetenue (pulteney 1799) thracia phaseolina (lamarck 1818) total for holocene skagen well: 71 (28.7%) geus bulletin no 3.pmd 28-06-2004, 08:4523 24 the danish late quaternary marine molluscs the record of shell-bearing danish late quaternary marine molluscs has been established on the basis of finds made during the systematic geological mapping since 1888, as presented in the publications from the geological survey of denmark, mainly in the i. række covering the descriptions for the map sheets. furthermore, special papers on holocene and late pleistocene marine molluscs have been included. most of them have been listed in one of the preceding chapters on works on danish sites with marine sediments. 1. the bælt sea area 2. the baltic sea area 3. the kattegat area with fjords 4. the limfjord area 5. the north sea coastal area 6. the vendsyssel area 7. the skagen area the regions are figured on the map (fig. 2) and follow the outline of the geological map sheets except the skagen area. it should be especially noted that the southern limit during the holocene of the north sea coastal region is at blåvands huk. although this area from varde to the german border has been mapped, no descriptions have been published. however, studies of the eemian from this region have been published and will be commented upon when presenting the eemian records from the above-mentioned areas. the bælt sea area has been taken as the region between the kattegat region and the baltic, here following ekman (1953) saying that the boundary between the bælt sea and the baltic proper is the threshold between gedser and darss and the southern end of the øresund, see fig. 1. the reason for including the øresund north of saltholm (northern and middle øresund sensu jensen & knudsen 1995) in the kattegat region is that regarding the water from the baltic the øresund is a sound – a passage – for the brackish water flowing north, but regarding the salt-water from the kattegat, the øresund is a fjord down to the threshold between amager and limhamn, to quote thorson (1944a, p. 42). the subfossil shell-bearing molluscs are presented on the background of the annotated check list of recent marine molluscs of danish waters by jensen & knudsen (1995). furthermore, general information has been taken from jensen & spärck (1934), lemche (1948), nordsieck (1968, 1969), fretter & graham (1976–1978, 1980–1982, 1984), fretter et al. (1986) and poppe & goto (1991, 1993) (which will not be quoted throughout the text, listing the many species and their environment). only the synonyms mentioned in texts on danish molluscan finds are included. a list of these synonyms is found heading the index of species. the investigations carried out by scientists in the northern atlantic have given a good base for the evaluation of the danish late quaternary molluscan fauna. these studies have been published mainly in papers on the zoology of east greenland, the godthaab expedition 1928, the zoology of iceland, and the zoology of the faroes. however, the zoology of greenland has been supplemented by contributions from macpherson (1971), lubinsky (1980), the 6. og 7. thule expedition til sydøstgrønland 1931–33 under the leadership of knud rasmussen and the treaarsexpeditionen til christian den x’s land 1931–34 under the leadership of lauge koch. in the two last-mentioned contributions, especially the animal ecology and the arctic communities have been treated, which form a very important part in the discussion of the danish late quaternary molluscan assemblages. the following publications on molluscs can be mentioned: 1. from greenland: spärck (1933), thorson (1933, 1944b, 1951), thorson & ussing (1934), madsen (1936), bertelsen (1937), lemche (1941a, b), ockelmann (1958), kramp (1961, 1963). 2. from iceland: spärck (1937), lemche (1938), thorson (1941), knudsen (1949a, b), madsen (1949). 3. from the faeroes: lemche (1928), thorson & spärck (1928), spärck & thorson (1931), petersen (1968), knudsen (1970). as to the community concept as worked out by c.g.j. petersen, references have already been given to petersen & jensen (1911), petersen (1913, 1914, 1915, 1918), and thorson (1957). the regional division of the northern european seas in to arctic, subarctic, boreal and lusitanian is from the zoogeographical division of the geus bulletin no 3.pmd 28-06-2004, 08:4524 25 northern european seas as given by feyling-hanssen (1955, fig. 5) and símonarson et al. (1998), as seen on figs 4 and 5 respectively. the molluscan genera and species are presented in groups within class, subclass and order mainly following the presentation of recent marine molluscs of danish waters as given by jensen & knudsen (1995). when there is no subfossil record at hand the information is taken from the recent data as given by poppe & goto (1991, 1993), jensen & knudsen (1995) and others as listed previously. to facilitate the use of the index of all molluscan species, a list of synonyms is given as mentioned earlier including the species mentioned in danish mollusc literature. class polyplacophora order neoloricata among the seven species recorded in the recent danish fauna only one, tonicella marmorea, has been found subfossil in the vendsyssel area from the younger weichselian deposits. it is a circumboreal species which in europe is known from northern scandinavia south to denmark and recorded from around iceland (knudsen 1949a, b) down to ireland (poppe & goto 1991). it is common in danish waters, being known from the central part of the kattegat and the øresund region (muus 1959), e.g. subarctic–boreal–lusitanian species. habitat. common in shallow water of less than 20 m, otherwise recorded from 0–183 m (muus 1959). the lack of information on subfossil finds may reflect difficulty in determination when the species is found in the subfossil state of preservation with the shell parts apart. polyplacophora do occur as seen from the skagen well where finds have been recognised at the 70.10–70.30 m and the 37.0–37.25 m levels, viz. during the subboreal and the subatlantic respectively. among the other – only recent – finds of polyplacophora, leptochiton asellus is widely found from the arctic to the lusitanian. hanleya hanleyi, ischnochiton albus and the above-mentioned tonicella marmorea are all found in the subarctic, boreal and lusitanian regions. the last three species, callochiton septemvalvis, lepidochitona cinereus and tonicella rubra, are all restricted to the boreal–lusitanian region. class gastropoda subclass prosobranchia order archaeogastropoda scissurella crispata fleming 1828 fig. 6 distribution. w greenland, s and w iceland, spitsbergen, norway north of lofoten, and south to the mediterranean. occurrence. the subarctic (not in true arctic water), boreal and lusitanian regions (according to thorson 1941). habitat. in the waters around iceland (thorson 1941, p. 4), the living specimens have often been found at depths greater than 500 m on clay bottom. however, poppe & goto (1991, p. 64) write that the species lives on stones, shelly sand and clay bottoms between 15 and 600 m. fretter & graham (1976, pp. 2–4) stated that the species is always sublittoral, even in the extreme northerly limits of its range, and occurs from 8– 2000 m. it is not recorded from danish waters although fretter & graham mentioned it from the norwegian and swedish coasts of the skagerrak. only subfossil finds. during the eemian in the kattegat region. fig. 6. scissurella crispata fleming 1828. anholt ii, 78.06–78.09 m b.s. × 20. mguh 25316. geus bulletin no 3.pmd 28-06-2004, 08:4525 26 patella vulgata linnaeus 1758 distribution. faeroes, norway off the lofoten islands, and south to the straits of gibraltar (thorson 1941). occurrence. boreal–lusitanian. habitat. intertidal, rocky shores or man-made hard substrates. however, it also occurs with seaweed. only one recent individual has been found at løkken, while empty shells are often found on the skagerrak coast (knudsen 1993). subfossil finds. the limfjord region, holocene, in a ‘køkkenmødding’ (kitchen midden) at bulbjerg (petersen 1888). helcion pellucidum (linnaeus 1758) distribution. west and south iceland, northern norway north of lofoten, and south to portugal (thorson 1941). also found in the øresund, but absent from the baltic and the limfjord. occurrence. boreal and lusitanian. habitat. lives on seaweeds at depths from 0 to 27 m. petersen (1888) points to this habitat as a reason why it is rarely found. subfossil finds. the limfjord region and vendsyssel, holocene. lepeta caeca (müller 1776) distribution. from northern scandinavia, n and s iceland and spitsbergen (thorson 1941) south to scotland. it is not present in the baltic, the north sea and the channel, but occurs in the azores. this species, according to poppe & goto (1991), prefers cold temperatures and lives at greater depths in the southern part of its range. in this way it exemplifies the tropical submerge. it is present in the øresund region and has been reported from single finds in the kattegat (petersen 1888). occurrence. subarctic and boreal. subfossil finds. none. iothia fulva (müller 1776) distribution. northern scandinavia and s and w iceland south to the irish sea, and like lepeta caeca in deep water off the azores. it is found in the øresund region, but with rare and single finds in the kattegat (petersen 1888). occurrence. boreal and lusitanian. habitat. offshore between 5 and 600 m on hard substrates like lepeta caeca. subfossil finds. the limfjord region, holocene. acmaea tessulata (müller 1776) distribution. east greenland, around iceland (thorson 1941) and in scandinavia according to petersen (1888). the limfjord (petersen 1986a). common in the øresund, but absent in the baltic. extends southwards to the north of the british isles and northern ireland. occurrence. arctic, subarctic, boreal, and lusitanian. habitat. lower part of the intertidal zone. in the south connected with zostera. subfossil finds. the limfjord region, holocene. acmaea virginea (müller 1776) distribution. northern scandinavia and around iceland (thorson 1941), south to the cape verde islands. common in the kattegat and øresund, but not found in the baltic. occurrence. subarctic, boreal, lusitanian. habitat. on hard substrate at depths from 0 to 100 m. subfossil finds. the limfjord and vendsyssel regions, holocene. emarginula fissura (linnaeus 1758) distribution. scandinavia and south to the mediterranean, according to petersen (1888) taken alive from the øresund, dead shells in the northern kattegat. geus bulletin no 3.pmd 28-06-2004, 08:4526 27 occurrence. boreal and lusitanian. habitat. from the low tide line to a depth of 700 m on hard substrate. subfossil finds. none. puncturella noachina (linnaeus 1771) distribution. spitsbergen and around iceland (thorson 1941), scandinavia south to portugal. also found in the skagerrak and kattegat including øresund. occurrence. arctic, subarctic, boreal and lusitanian. habitat. between 10 and 200 m on rock and stones. subfossil finds. none. margarites helicinus (phipps 1774) distribution. from northern scandinavia, spitsbergen and around iceland (thorson 1951) south to the british isles. a few records from the skagerrak and kattegat. occurrence. arctic, subarctic, boreal and northern part of the lusitanian region. habitat. from the intertidal zone to 400 m deep on seaweeds and under stones. subfossil finds. the limfjord region, holocene. gibbula cineraria (linnaeus 1758) distribution. from northern scandinavia north of lofoten, and w and s iceland (thorson 1941) south to morocco. according to petersen (1888) known from the kattegat including øresund, the bælt sea and the limfjord area. occurrence. boreal–lusitanian. habitat. intertidal to 130 m deep on rocks and seaweeds. subfossil finds. the limfjord, the north sea and vendsyssel, holocene. recorded from the north sea during the eemian. gibbula tumida (montagu 1803) distribution. from northern norway north of lofoten, and s and w iceland (thorson 1941), south to spain. known from the kattegat, including the øresund, but not so common as gibbula cineraria. the species does not occur in the limfjord (petersen 1986a, table 1). occurrence. boreal–lusitanian. habitat. on gravel bottoms from below low tide to depths of 1200 m. subfossil finds. the limfjord and vendsyssel area, holocene. jujubinus clelandi (w. wood 1828) distribution. from the lofoten islands south into the mediterranean. found in the kattegat region, including the øresund, but rare. occurrence. boreal–lusitanian. habitat. on various types of bottom from depths of 35 m to 800 m. subfossil finds. none. calliostoma formosa (mighels 1842) distribution. along the norwegian coast, w and s iceland (thorson 1941), and south to the british isles but not on the west side. occurrence. boreal. habitat. dredged at depths between 19 and 1000 m – the species is never littoral. subfossil finds. none. calliostoma zizyphinum (linnaeus 1758) distribution. from the lofoten islands south to the azores. in the skagerrak. occurrence. boreal–lusitanian. geus bulletin no 3.pmd 28-06-2004, 08:4527 28 habitat. intertidal to 300 m deep – lives on all types of bottoms. subfossil finds. none. skenea serpuloides (montagu 1808) distribution. from the british isles south to portugal and into the mediterranean. occurrence. lusitanian. habitat. from the intertidal zone down to 50 m deep. intertidal on weeds and stones, and sublittoral dredged from shelly and gravelly sand. only subfossil finds. the limfjord area, holocene. skenea basistriata (jeffreys 1877) distribution. atlantic coast of europe, but not in the north sea and the baltic. occurrence. boreal–lusitanian. habitat. on soft bottom – never in shallow water or littoral sequences, deep water 90–2400 m. subfossil finds. the limfjord region, holocene. theodoxus fluviatilis (linnaeus 1758) distribution. from the pyrenees and the british isles east towards the caucasus, including northern sweden and the coasts of finland. occurrence. lusitanian–boreal. habitat. the primary habitat of this species is rivers. in the baltic sea the form littoralis becomes a common littoral animal (fretter & graham 1978a, p. 105). the species is also recorded from the fjords bordering the kattegat region today. subfossil finds. the kattegat and limfjord regions, holocene. the archaeogastropoda are represented by 16 species in the recent marine fauna, out of which ten have been recorded from the past. only two species, skenea serpuloides and scissurella crispata, are not found in the recent fauna. there is a total number of subfossil finds among the archaeogastropoda of 12 species, with two from the eemian and ten from the holocene. order mesogastropoda littorina littorea (linnaeus 1758) distribution. from northern norway north of lofoten, and south to spain. common along all the danish coasts but not in the baltic and on the more exposed sandy coasts (petersen 1888). recent records from the baltic as far as bornholm (fretter & graham 1980, p. 256). occurrence. boreal and lusitanian. habitat. the intertidal zone, abundant on rocky shores, might be found to a depth of 60 m. subfossil finds. the bælt sea, baltic, kattegat, limfjord, the north sea and vendsyssel regions, holocene. the bælt sea, baltic, kattegat and north sea regions in the eemian. melaraphe (littorina) neritoides (linnaeus 1758) distribution. from western norway south to morocco and the mediterranean. a scattered occurrence in the kattegat (jensen & knudsen 1995). the species is recorded from the limfjord (petersen 1986a). occurrence. boreal and lusitanian. habitat. lives high on the rocky shores (splash zone). subfossil finds. none. littorina mariae sacchi & rastelli 1966 distribution. from northern scandinavia south to the mediterranean, extending through the kattegat into the bælt sea. occurrence. the boreal and lusitanian. habitat. in the tidal zone on weeds. subfossil finds. none. (the species may have been confused with littorina obtusata.) geus bulletin no 3.pmd 28-06-2004, 08:4528 29 littorina obtusata (linnaeus 1758) distribution. w greenland, s and w iceland (thorson 1941), northern scandinavia north of lofoten, and south to the mediterranean. the species extends through the limfjord and kattegat into the bælt sea (fretter & graham 1980). occurrence. subarctic, boreal and lusitanian. habitat. intertidal, lives on weeds. subfossil finds. the bælt sea, kattegat, limfjord, north sea and vendsyssel regions, holocene. during the eemian recorded from the north sea. littorina saxatilis (olivi 1792) distribution. from greenland, spitsbergen, around iceland and the atlantic coasts of europe. common in the fjords bordering the kattegat, including the limfjord (petersen 1986a). occurrence. the arctic, subarctic, boreal and lusitanian. habitat. intertidal. subfossil finds. the bælt sea, kattegat, limfjord, north sea and vendsyssel regions, holocene. the north sea during the eemian, and the vendsyssel area in the late weichselian. littorina tenebrosa (montagu 1803) distribution. off southern iceland and the atlantic coasts of europe. the species is found in the limfjord, and penetrates also into the baltic, with finds off møn and stevns (petersen 1888). occurrence. the boreal and lusitanian. habitat. intertidal. subfossil finds. the bælt sea, baltic, kattegat, limfjord and vendsyssel regions, holocene. lacuna pallidula (da costa 1778) fig. 7a, b distribution. from spitsbergen, around iceland and along the atlantic coast down to the gulf of biscay. enters the danish waters, including the limfjord, the øresund and the bælt sea. occurrence. the arctic, subarctic, boreal and lusitanian. habitat. intertidal to 70 m deep on weeds. subfossil finds. the bælt sea, limfjord, north sea and vendsyssel regions, holocene. in the skagen well from the subatlantic. lacuna crassior (montagu 1803) distribution. from the arctic seas to the british isles. only records from the north sea and from the nw coast of sweden. occurrence. the arctic, subarctic and boreal. habitat. sublittoral to 90 m deep on soft bottoms with stones and shells (fretter & graham 1980, p. 248). subfossil finds. none. fig. 7. a, b: lacuna pallidula (da costa 1778). skagen 4, 30.8– 30.5 m b.s., lab. no. 355,93. × 20. mguh 25317. geus bulletin no 3.pmd 28-06-2004, 08:4529 30 lacuna parva (montagu 1803) distribution. from norway off the lofoten, and south to spain, found in the kattegat region with fjords, but few records, and not found in the limfjord. occurrence. boreal and lusitanian. habitat. intertidal extending sublittorally to around 50 m as lacuna vincta (fretter & graham 1980, p. 250) and living on seaweeds. subfossil finds. the limfjord, north sea and vendsyssel regions, holocene. from the north sea during the eemian. lacuna vincta (montagu 1803) distribution. from w greenland around iceland, norway and south to spain. in the danish waters, including the limfjord, found into the bælt sea (fretter & graham 1980). occurrence. subarctic, boreal and lusitanian. habitat. intertidal to 60 m deep living on seaweeds. subfossil finds. the bælt sea, kattegat, limfjord, north sea and vendsyssel regions, holocene. in the north sea region found during the eemian. in the vendsyssel area recorded from the eemian, and also from the early/middle and late weichselian (older and younger yoldia sea respectively). hydrobia neglecta muus 1963 distribution. the british isles, ireland and the north sea. occurrence. boreal and lusitanian. habitat. shallow-water environments on the soft substratum or the vegetation. subfossil finds. none. hydrobia ulvae (pennant 1777) fig. 8 distribution. norway off lofoten south to the mediterranean. in all the danish waters including the baltic. occurrence. the boreal and lusitanian. habitat. the intertidal zone, but has been found as deep as 20 m, on soft substrate, most often on intertidal banks of firm mud or muddy sand (fretter & graham 1978a, p. 122). subfossil finds. the bælt sea, baltic, kattegat, limfjord, north sea and vendsyssel regions, holocene. in the skagen well from the subatlantic. during the eemian recorded from the bælt sea, baltic, kattegat and the north sea regions. hydrobia ventrosa (montagu 1803) distribution. norway off lofoten, south to the mediterranean. in all the danish waters including the baltic. occurrence. the boreal and lusitanian regions. habitat. the intertidal zone,on soft substratum like hydrobia ulvae, but prefers lower salinities (fretter & graham 1978a, p. 126). the two species may occur together, so quantitative analyses must be undertaken to designate any changes in the environment (petersen 1993). subfossil finds. the baltic, kattegat, limfjord and the north sea regions, holocene. potamopyrgus antipodarum (gray 1853) distribution. from scandinavia to spain and in danish waters into the baltic, a late immigrant according to jensen & knudsen (1995). occurrence. the boreal and lusitanian regions. fig. 8. hydrobia ulvae (pennant 1777). skagen 4, 25.0–25.5 m b.s., lab. no. 350,93. × 20. mguh 25318. geus bulletin no 3.pmd 28-06-2004, 08:4530 31 habitats. in all kinds of brackish and freshwater habitats. much like the distribution of hydrobia ventrosa in brackish waters (fretter & graham 1978a, p. 132). subfossil finds. none. skeneopsis planorbis (fabricius 1780) distribution. w greenland, around iceland and norway, south to the mediterranean. known from few places in danish waters. occurrence. the subarctic, boreal and lusitanian regions. habitat. intertidal down to a depth of 70 m, lives on seaweeds. only subfossil finds. the limfjord and the vendsyssel regions, holocene. barleeia unifasciata (montagu 1803) fig. 9 distribution. from the shetlands south into the mediterranean. occurrence. the boreal and lusitanian regions. habitat. shallow waters, intertidal, lives on seaweeds on rocky shores. only subfossil finds. the skagen well from the subatlantic. alvania abyssicola (forbes 1850) distribution. from northern norway to the mediterranean. the species extends into the skagerrak and kattegat, including the øresund. occurrence. the boreal and lusitanian regions. habitat. on muddy bottom in sublittoral areas at depths of 15–100 m. subfossil finds. the vendsyssel area from the eemian. alvania jeffreysi (waller 1864) distribution. s and w iceland, norway, south to the mediterranean. found in the skagerrak, but not in the north sea. occurrence. the boreal and lusitanian regions. habitat. always sublittoral from 50 to 600 m on sandy bottom. subfossil finds. none. alvania lactea (michaud 1830) distribution. from the channel islands south to morocco and the mediterranean. occurrence. the lusitanian region. habitat. sublittorally under stones and amongst algae. only subfossil finds. the limfjord and vendsyssel areas, holocene. alvania cimicoides (forbes 1844) distribution. sw and nw iceland (empty shells), norway, north of lofoten, and south to the mediterranean; probably absent from the channel and the north sea. occurrence. the boreal and lusitanian regions. habitat. sublittoral, from the laminarian zone downwards, but mainly in deeper water down to 500 m. fig. 9. barleeia unifasciata (montagu 1803). skagen 3, 33.90– 34.20 m b.s., lab. no. 709,93. × 20. mguh 25319. geus bulletin no 3.pmd 28-06-2004, 08:4531 32 usually found on soft bottoms (fretter & graham 1978b). only subfossil finds. the vendsyssel area, holocene. alvania punctura (montagu 1803) distribution. norway south of lofoten (including lofoten) and south to the mediterranean. it extends to the swedish west coast, but is absent from the øresund, the baltic, the eastern shores of the north sea and the eastern basin of the channel (fretter & graham 1978b). however, according to jensen & knudsen (1995) alvania punctura is found in the northern and central parts of the øresund. occurrence. the boreal and lusitanian regions. habitat. sublittoral to depths of c. 100 m, on both finer and coarser substrata. subfossil finds. the limfjord and the vendsyssel regions, holocene. alvania cruenta odhner 1915 distribution. arctic canada, west greenland and svalbard (thorson 1951; macpherson 1971). occurrence. the arctic and subarctic. habitat. from 19 to 234–254 m on mud (macpherson 1971). only subfossil finds. the early/middle weichselian in the vendsyssel region. alvania jan mayeni (friele 1886) distribution. e and w greenland, spitsbergen, ne iceland, and norway north of lofoten (thorson 1941). occurrence. the arctic and subarctic with boreal outposts. habitat. off iceland between 94–442 m in deep on clay with many stones (thorson 1941). only subfossil finds. the early/middle weichselian (older yoldia clay) in vendsyssel. alvania scrobiculata (möller 1842) distribution. e and w greenland, spitsbergen, n and e iceland, and norway north of lofoten (thorson 1941). occurrence. the arctic and subarctic with boreal outposts. habitat. from 22 m at e greenland to 342 m in the northern arctic sea on a bottom of sand and with algae (thorson 1941). only subfossil finds. from the early/middle weichselian (older yoldia clay) in vendsyssel. cingula semistriata (montagu 1808) distribution. from lofoten and south along the west coast of norway to the mediterranean (rare in the north). it extends into the kattegat, but is absent from the eastern shores of the southern north sea and the limfjord. occurrence. the boreal and lusitanian regions. habitat. on rocky shores in the intertidal zone and sublittorally to 100 m, fond of silty places (fretter & graham 1978b). subfossil finds. the limfjord and the vendsyssel areas, holocene. cingula turgida (jeffreys 1870) distribution. from norway north of lofoten to the south into kattegat. occurrence. the boreal region. habitat. on muddy bottoms down to c. 1000 m. subfossil finds. the north sea region, holocene. obtusella alderi (jeffreys 1858) distribution. from norway to spain, not on the eastern shores of the north sea and in the baltic; however, found in the kattegat, including the northern part of the øresund. occurrence. the boreal and lusitanian regions. geus bulletin no 3.pmd 28-06-2004, 08:4532 33 habitat. sublittoral to a depth of 60 m amongst algae and on sandy or gravelly bottoms. subfossil finds. none. onoba aculeus (gould 1841) distribution. from spitsbergen, w greenland, around iceland and along the coast of norway into the kattegat, including the øresund. also found at localities off ireland. occurrence. the arctic, subarctic, boreal and lusitanian regions. habitat. found to about 200 m on algae. subfossil finds. none. onoba semicostata (montagu 1803) distribution. around iceland, along the coast of norway and south to the mediterranean. absent from the eastern north sea coasts, but extends through the limfjord and kattegat, including the øresund, into the bælt sea and the most saline parts of the baltic (fretter & graham 1978b). occurrence. the subarctic, boreal and lusitanian regions. habitat. from the intertidal region to depths of 100 m. found under stones, amongst weeds, mussels and tunicates in shelly gravel, but only where there are quantities of silt (fretter & graham 1978b). subfossil finds. the bælt sea, kattegat, limfjord and vendsyssel regions, holocene. onoba proxima (forbes & hanley 1850) distribution. from the western coast of britain south to the mediterranean. occurrence. the lusitanian region. habitat. from 10 to 170 m on bottoms of muddy sand. only subfossil finds. the limfjord region, holocene. onoba vitrea (montagu 1803) fig. 10 distribution. from norway off lofoten and south to the mediterranean, extends through skagerrak into the kattegat, including the øresund, but absent from the limfjord. occurrence. the boreal and lusitanian regions. habitat. muddy bottoms at depths of 10–50 m in the northern parts of its range, but extending to 120 m in the south. further notes on onoba species in fretter & graham (1978b, p. 170). subfossil finds. the kattegat, limfjord, north sea, vendsyssel and skagen areas, holocene. in the skagen well recorded from the atlantic, the subboreal and the subatlantic. during the eemian recorded from the north sea. rissoa albella lovén 1846 fig. 11 distribution. from norway off the lofoten islands, and south to the mediterranean, extending into the limfjord and the kattegat with the danish fjords, including the øresund and the bælt sea. occurrence. the boreal and lusitanian regions. habitat. on rocky shores amongst weeds, sublittoral to 15 m. it is tolerant of some brackishness (fretter & graham 1978b). subfossil finds. the bælt sea, baltic, kattegat, limfjord, north sea, vendsyssel and skagen well from the subfig. 10. onoba vitrea (montagu 1803). skagen 3, 67.0–67.25 m b.s., lab. no. 720,93. × 9.6. mguh 25320. geus bulletin no 3.pmd 28-06-2004, 08:4533 34 atlantic, holocene. from the eemian recorded from the north sea. rissoa inconspicua alder 1844 distribution. from northern norway, north of lofoten, and south to the mediterranean. occurring in the limfjord, øresund and bælt sea. occurrence. the boreal and lusitanian regions. fretter & graham (1978b, p. 200) indicate that the species is found to the arctic. however, it is not recorded from iceland and the faroes but only from norway north of lofoten (thorson 1941, table ii, p. 141). habitat. typically sublittoral living on algae, and on sandy gravel to depths of about 100 m. tolerant of slightly brackish conditions. subfossil finds. the bælt sea, baltic, kattegat, limfjord, north sea and vendsyssel, holocene. during the eemian recorded from the bælt sea, kattegat and north sea regions. rissoa membranacea (j. adams 1800) distribution. from norway off lofoten and south to the canary islands. extends into the limfjord, kattegat, bælt sea and the westernmost part of the baltic (the rügen island). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone sublittorally to about 15 m associated with zostera or on weeds with the same habit, extending into brackish water. subfossil finds. the bælt sea, baltic, kattegat, limfjord, north sea and vendsyssel areas, holocene. during the eemian in the bælt sea and north sea regions. rissoa parva (da costa 1779) distribution. from norway, north of lofoten and the faeroes, south to the mediterranean, found in the limfjord (petersen 1986a) and the northern part of the øresund but not in the baltic according to fretter & graham (1978b). however, bondesen (1975) includes the species in the baltic, but excludes it from the bælt sea. petersen (1888, p. 93) regarded the species as being limited to the central part of the kattegat, depending on sufficient salt content – therefore the information of the occurrences in the baltic given by bondesen (1975) is surprising. occurrence. the boreal and lusitanian regions. habitat. from the tidal zone to about 25 m on fronds, smaller weeds and under stones. in the faeroes from the rock pools and the beach to a depth of 20 m (spärck & thorson 1931). subfossil finds. the limfjord and vendsyssel areas, holocene. during the eemian recorded from the bælt sea, north sea and vendsyssel regions. rissoa violacea desmarest 1814 fig. 12 distribution. rissoa violacea is not at all recorded from iceland (thorson 1941) and the faroes (spärck & thorson 1931) but from norway off the lofoten islands and south. both of the recent subspecies are recorded from the kattegat, including the øresund region. bonfig. 12. rissoa violacea desmarest 1814. skagen 4, 27.0– 27.5 m b.s., lab. no. 352,93. × 20. mguh 25322. fig. 11. rissoa albella lovén 1846. skagen 4, 30.0–30.5 m b.s., lab. no. 355,93. × 20. mguh 25321. geus bulletin no 3.pmd 28-06-2004, 08:4534 35 desen (1975) has rissoa violacea sensu lato from the skagerrak, the kattegat and the limfjord regions. occurrence. the boreal and lusitanian regions. habitat. in the tidal zone to about 50 m on weeds and amongst sandy gravel. subfossil finds. the limfjord, north sea, vendsyssel and skagen areas, holocene. in skagen recorded from the subatlantic. the species occurred during the eemian in the kattegat region. assiminea grayana fleming 1828 distribution. the species is confined to the north sea coasts. indanishwaters it extendssouth toblåvandshuk. occurrence. the boreal region. habitat. the species is limited to the upper parts of the salt-marsh areas on the vegetation. subfossil finds. none. caecum glabrum (montagu 1803) distribution. from norway off lofoten south to the mediterranean. it extends into kattegat, where it has been reported from the northern part (jensen & knudsen 1995), but not from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. sublittorally to about 250 m on sandy and sandy–muddy bottoms. subfossil finds. the limfjord, north sea and vendsyssel regions, holocene. recorded from the bælt sea and the north sea during the eemian. tornus exquisitus (jeffreys 1883) distribution. within the danish waters recorded from northern kattegat (jensen & knudsen 1995), but no record is given by fretter & graham (1978b, p. 232) so the species is not treated further. habitat. unknown for this species. subfossil finds. none. bittium reticulatum (da costa 1778) fig. 13 distribution. from off lofoten in norway south to the mediterranean. it is found through the skagerrak, limfjord, kattegat, including the øresund, and into the bælt sea, but not from the southern coastal part of the north sea. occurrence. the boreal and lusitanian regions. habitat. common in shallow sublittoral water, but recorded to 250 m. found on soft bottoms in association with weeds. subfossil finds. the species is recorded from all regions during the holocene. from skagen recorded from the subatlantic. from the eemian found in the bælt sea, kattegat, north sea, and vendsyssel regions. turritella communis risso 1826 fig. 14 distribution. from the lofoten islands south to the mediterranean. the species extends into all the danish waters as far as the øresund. occurrence. the boreal and lusitanian regions. habitat. from 10 to 200 m depths on soft bottoms. subfossil finds. the limfjord, north sea, vendsyssel and skagen areas, holocene. in the skagen area refig. 13. bittium reticulatum (da costa 1778). geus collection. vejlager xiv, denmark. × 9.6. mguh 25323. geus bulletin no 3.pmd 28-06-2004, 08:4535 36 corded from the subboreal and subatlantic. during the eemian recorded from the baltic, kattegat, north sea and vendsyssel regions. turritella erosa couthouy 1838 distribution. west and east greenland (thorson 1944b, p. 40), svalbard and on the northern coast of russia (macpherson 1971). occurrence. the arctic and subarctic regions. habitat. from 10 to 350 m on soft bottoms. subfossil finds. recorded from the vendsyssel region in early/middle weichselian. aporrhais pespelicani (linnaeus 1758) fig. 15 distribution. on the southern and western part of iceland, norway off lofoten with a questionable occurrence north of lofoten (thorson 1941), and south to the mediterranean. it does not occur off the west coast of denmark nor in its fjords, except the limfjord. it extends through the kattegat, including øresund (fretter & graham 1981, p. 297). empty shells are found in kieler bucht (arntz et al. 1976). occurrence. the boreal and lusitanian regions. habitat. sublittoral to depths of 180 m on mud, muddy sand and sand. subfossil finds. the baltic, limfjord, north sea, vendsyssel and skagen areas, holocene. in the skagen region from the subatlantic. during the eemian recorded from the bælt sea, baltic, north sea, and vendsyssel regions. aporrhais serresianus (michaud 1828) distribution. from southern and western iceland, norway off lofoten and south to the mediterranean. according to fretter & graham (1981) absent from the skagerrakand all danish seas; however, jensen & knudsen (1995) note a single record from the central kattegat. occurrence. the boreal and lusitanian regions. habitat. from sublittoral (as aporrhais pespelicani) down to 1000 m on finer muds. subfossil finds. none. crepidula fornicata (linnaeus 1758) distribution. c. fornicata is a late immigrant to europe (first recorded in the british isles late in the 19th century). the present distribution is from norway south to portugal. it reached the limfjord in 1934. in 1949– 50 it was found in the northern part of kattegat but has not spread further south (jensen & knudsen 1995). fig. 15. aporrhais pespelicani (linnaeus 1758). skagen 3, 0.0–30.0 m b.s. (washed sample). × 4.8. mguh 25325. fig. 14. turritella communis risso 1826. skagen 3, 73.10–73.30 m b.s., lab. no. 722,93. × 4.8. mguh 25324. geus bulletin no 3.pmd 28-06-2004, 08:4536 37 occurrence. the boreal–lusitanian regions within its european distribution. crepidula fornicata was transferred by man as seen also for species like mya arenaria. habitat. sublittoral to depths of c. 10 m. the animals live in chains, the oldest attached to a substrate which might be an oyster. the species was actually transported to europe with oysters (fretter & graham 1981, p. 311). subfossil finds. obviously none. capulus ungaricus (linnaeus 1758) distribution. empty shells recorded from sw and nw iceland and living specimens from norway north of lofoten south to the mediterranean. the record from greenland mentioned in fretter & graham (1981) cannot be sustained in the literature (thorson 1944b, 1951). occurrence. the boreal and lusitanian regions. habitat. usually sublittorally to 805 m attached to stones or the host animal. subfossil finds. none (young date from the north sea). lamellaria perspicua (linnaeus 1758) distribution. sw and nw iceland, norway from lofoten and south to the mediterranean. it occurs in the skagerrak but not on the danish coasts. occurrence. the boreal and lusitanian regions. habitat. from the tidal zone and downwards to depths of 1200 m, especially in the southern parts of its range on rocky shores and under stones. subfossil finds. none. velutina plicatilis (müller 1776) distribution. from east greenland and spitsbergen, southern and western iceland, norway north of lofoten and south to northern spain. it extends into the skagerrak, kattegat and the northern part of the north sea. occurrence. the arctic, subarctic, boreal and lusitanian regions. habitat. from 10 to c. 375 m deep on hard bottoms, usually in association with ascidians and hydroids. subfossil finds. none. velutina velutina (müller 1776) distribution. from spitsbergen, e and w greenland, around iceland and norway south to the mediterranean. it extends into the skagerrak and kattegat, including øresund. occurrence. arctic, subarctic, boreal and lusitanian regions. habitat. sublittoral extending to 1000 m on hard bottoms associated with tunicates. subfossil finds. none. trivia arctica (pulteney 1799) distribution. from norway off lofoten and south to the mediterranean. it extends into the skagerrak and kattegat (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. sublittoral to about 100 m, in southerly latitudes to about 1000 m. it is associated with ascidians. subfossil finds. none. trivia monacha (da costa 1778) distribution. from the british isles south to the mediterranean. the recent distribution in the north sea is questioned, and the species is not recorded from scandinavia (fretter & graham 1981, p. 329). occurrence. the lusitanian region. habitat. on rocky shores and under stones associated with ascidians. subfossil finds. the vendsyssel region, holocene. geus bulletin no 3.pmd 28-06-2004, 08:4537 38 amauropsis islandicus (gmelin 1791) distribution. from spitsbergen, e and w greenland, around iceland and norway. it extends into the skagerrak and kattegat, including øresund, although it is rare there (jensen & knudsen 1995). occurrence. the arctic, subarctic and boreal regions. habitat. sublittorally to about 80 m deep on sandy clay bottoms. subfossil finds. none (late date from the north sea). lunatia alderi (forbes 1838) fig. 16 distribution. from southern and western iceland and norway off the lofoten islands south to the mediterranean (thorson 1941). it extends into the skagerrak, kattegat and øresund, but the species is not recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. sublittoral between 10 and 50 m, extending to 2000 m. infaunal on sandy shores, clean sand and some admixture of mud. according to petersen (1888), it is common on mixed bottoms in the kattegat. subfossil finds.the limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen area recorded from the atlantic, subboreal and subatlantic. from the eemian recorded from the baltic, kattegat, north sea, vendsyssel and skagen regions. lunatia catena (da costa 1778) distribution. from the skagerrak and kattegat, including øresund but not the limfjord, south to the mediterranean (fretter & graham 1981, p. 339). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone down to about 125 m on sandy bottoms. subfossil finds. the limfjord, north sea and vendsyssel regions, holocene. lunatia montagui (forbes 1838) fig. 17 distribution. from w and s iceland, norway north of the lofoten islands, and south to the mediterranean. it occurs in the skagerrak and kattegat, including øresund, however, rare in the sound as the presiding species according to jensen & knudsen (1995). occurrence. the boreal and lusitanian regions. habitat. from 15 to 200 m depth on sandy and muddy bottoms. subfossil finds. recorded from the skagen well from the subatlantic. fig. 16. lunatia alderi (forbes 1838). skagen 4, 26.0–26.5 m b.s., lab. no. 351,93. × 20. mguh 25326. fig. 17. lunatia montagui (forbes 1838). skagen 4, 20.0–20.5 m b.s., lab. no. 345,93. × 4.8. mguh 25327. geus bulletin no 3.pmd 28-06-2004, 08:4538 39 lunatia pallida (broderip & sowerby 1829) distribution. from spitsbergen, e and w greenland, around iceland and norway, south to the north sea, skagerrak and the kattegat, including the øresund. occurrence. the arctic, subarctic and boreal regions. habitat. from 10 to 2000 m on clay bottoms – the greatest depths in the most southerly parts of its range (fretter & graham 1981). in greenland waters commonin the arctic macoma community (thorson 1944b). subfossil finds. the vendsyssel region during the early/ middle weichselian and late weichselian (the older and younger yoldia sea respectively). natica affinis (gmelin 1790) distribution. from spitsbergen, w and e greenland, around iceland and norway south to the mediterranean. however, within the lusitanian region the species lives in deep water (fretter & graham 1981, p. 345). the species is recorded from danish waters (jensen & knudsen 1995). occurrence. the arctic, subarctic, boreal and lusitanian regions. habitat. from about 4 m depth in high latitudes to well over 2000 m in low ones on sandy, muddy and clay bottoms. as mentioned for other species with a wide geographical distribution this is a ‘tropical submerge’ which is quite common for cold-water animals which in the northern regions inhabit the surface water to occur mainly or exclusively in deeper zones in the southern seas (ekman 1953, p. 112). the species is found in the arctic macoma community (spärck 1937). subfossil finds. recorded from the vendsyssel area during the early/middle weichselian (the older yoldia sea), and the late weichselian (the younger yoldia sea). order heterogastropoda triphora adversa (montagu 1803) distribution. from norway off the lofoten islands and south to spain. it extends into the kattegat, including øresund and the bælt sea, but not recorded from the limfjord. occurrence. the boreal and lusitanian regions. habitat. sublittorally to 100 m under stones, algae or associated with sponges. subfossil finds. the bælt sea, kattegat, limfjord, north sea and vendsyssel areas, holocene. recorded from the bælt sea and the north sea during the eemian. cerithiella metula (lovén 1846) distribution. s and w iceland, norway north of lofoten, and south to the mediterranean. recorded from the danish waters (jensen & knudsen 1995), although rare in the skagerrak and not occurring in the kattegat (fretter & graham 1982, p. 377). occurrence. the boreal and lusitanian regions. habitat. from 40 to 400 m depth on soft bottoms. subfossil finds. none. cerithiopsis barleei jeffreys 1867 distribution. according to fretter & graham (1982), from sw england south to the mediterranean. however, the species is recorded from danish waters, i.e. the øresund area, although as rare (jensen & knudsen 1995), but is not found in the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. sublittorally associated with sponges. subfossil finds. the limfjord region, holocene. geus bulletin no 3.pmd 28-06-2004, 08:4539 40 cerithiopsis tubercularis (montagu 1803) distribution. from norway off the lofoten islands and south to the mediterranean. recorded from the swedish west coast, but not in danish waters (fretter & graham 1982). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone and sublittorally to 100 m. the species is found on sponges. only subfossil finds. the limfjord region, holocene, and the north sea during the eemian. epitonium clathratulum (kanmacher 1797) distribution. from norway and south to the mediterranean. it extends into the kattegat, including the øresund. occurrence. the boreal and lusitanian regions. habitat. from 30 to 100 m deep on sandy–muddy bottoms. subfossil finds. none. epitonium clathrus (linnaeus 1758) distribution. from norway off the lofoten islands and south to the mediterranean. the species extends into the kattegat and øresund (jensen & knudsen 1995), but does not enter the danish fjords (fretter & graham 1982, p. 387). occurrence. the boreal and lusitanian regions. habitat. sublittorally from 5 to 70 m on sandy–muddy bottoms. subfossil finds. the limfjord, north sea and vendsyssel regions, holocene. recorded from the bælt sea and the north sea during the eemian. epitonium trevelyanum (johnston 1841) fig. 18 distribution. from norway off the lofoten islands and south to the mediterranean. the species extends into the kattegat, including the øresund. occurrence. the boreal and lusitanian regions. habitat. from 30 to 200 m depth on sandy–muddy bottoms. subfossil finds. the skagen area, holocene, recorded from the subboreal and the subatlantic. epitonium turtonis (turton 1819) distribution. from norway off the lofoten islands (thorson 1941) and south to the mediterranean. occurring in the kattegat (fretter & graham 1982) and øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from 5 to 20 m deep on sandy–muddy bottoms. subfossil finds. the limfjord and vendsyssel regions, holocene. aclis ascaris (turton 1819) distribution. from norway off lofoten and south to the mediterranean. the species extends into the skagerrak and kattegat, including the øresund. occurrence. the boreal and lusitanian regions. habitat. from 10 to 50 m deep on soft sandy bottoms. subfossil finds. the north sea, holocene. fig. 18. epitonium trevelyanum (johnston 1841). skagen 3, 58.5– 60.0 m b.s., lab. no. 98,93. × 9.6. mguh 25328. geus bulletin no 3.pmd 28-06-2004, 08:4540 41 aclis minor (brown 1827) fig. 19 distribution. from norway off the lofoten islands and south to the mediterranean. the species is recorded from the southern kattegat and øresund. occurrence. the boreal and lusitanian regions. habitat. from 15 to 150 m deep on bottoms of sand, muddy sand or gravel. subfossil finds. the limfjord, north sea and skagen areas, holocene; in the skagen well recorded from the atlantic, subboreal and subatlantic. aclis walleri jeffreys 1867 distribution. from norway off the lofoten islands and south to the mediterranean. recorded from north of skagen (petersen 1888) and the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. according to fretter & graham (1982), at greater depths than the other aclis ssp. – down to 550 m on soft bottoms. subfossil finds. the north sea, holocene. eulima bilineata (alder 1848) distribution. from norway north of lofoten, south to the mediterranean. in the danish waters recorded from the southern kattegat. occurrence. the boreal and lusitanian regions. habitat. from 20 to 250 m deep on soft bottoms associated with ophiuroids. subfossil finds. none. haliella stenostoma (jeffreys 1858) distribution. off west greenland, around iceland, norway off the lofoten islands, and south to the mediterranean. occurrence. the subarctic, boreal and lusitanian regions. habitat. sublittoral from about 70 to 3000 m on soft bottoms. the species has its main occurrence at rather great depths (thorson 1941), but is also recorded from danish waters (jensen & knudsen 1995). subfossil finds. none. polygireulima sinuosa (sacco 1836) fig. 20 distribution. from the kattegat and south to the mediterranean. occurrence. the boreal and lusitanian regions. habitat. sublittoral from 30 to 150 m on soft bottoms. subfossil finds. the skagen area, holocene, recorded from the subatlantic. fig. 19. aclis minor (brown 1827). skagen 3, 55.1–55.3 m b.s., lab. no. 716,93. × 20. mguh 25329. fig. 20. polygireulima sinuosa (sacco 1836). skagen 3, 38.19–38.24 m b.s., core sample k-6. × 9.6. mguh 25330. geus bulletin no 3.pmd 28-06-2004, 08:4541 42 polygireulima monterosatoi (monterosato 1890) distribution. norway south of the lofoten islands, and south to the mediterranean. occurrence. the boreal and lusitanian regions. habitat. from 20 to 120 m deep on sandy muddy or gravelly bottoms. however, as fretter & graham (1982, p. 421) say: “presumably these animals attack echinoderms like their relatives, but which is not known”. subfossil finds. none. vitreolina collensi (sykes 1903) fig. 21 distribution. from the west coast of britain and ireland to the mediterranean. fig. 21. vitreolina collensi (sykes 1903). skagen 3, 70.10–70.30 m b.s., lab. no. 721,93. × 20. mguh 25331. fig. 22. a, b: vitreolina philippii (rayneval & pouzi 1854). skagen 4, 21.0–21.5 m b.s., lab. no. 346,93. × 20. mguh 25332. occurrence. the lusitanian region. habitat. sublittorally to 35–40 m on soft bottoms. only subfossil finds. the skagen area, holocene, recorded from the subboreal and the subatlantic. vitreolina philippii (reyneval & ponzi 1854) fig. 22a, b distribution. from norway off the lofoten islands and south to the mediterranean. according to petersen (1888), very common in danish waters from the skagerrak, the kattegat and the øresund, and the bælt sea, but not recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from the lowest part of the tidal zone to a depthof 200 m on soft bottoms. fretter & graham (1982, p. 422) hold it as perhaps the most common local eulimid and like other eulimids an intermittent parasite of echinoderms. subfossil finds. the limfjord, vendsyssel and skagen areas, holocene, recorded from the subatlantic in the skagen well. occurring in the eemian in the vendsyssel region. graphis albida (kanmacher 1798) fig. 23 distribution. from southern norway south to the mediterranean. the species has a record from the limfjord (petersen 1986a). geus bulletin no 3.pmd 28-06-2004, 08:4542 43 occurrence. the boreal and lusitanian regions. habitat. from low in the tidal zone to 30 m deep on muddy and sandy bottoms. only subfossil finds. the skagen well area, holocene, recorded from the subatlantic. melanella lubrica (monterosato 1891) fig. 24 distribution. from norway south to iberia. the species is recorded from the danish waters south into the øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from 14 to 100 m on soft bottoms of muddy sand and gravel. the species is an intermittent ectoparasite (fretter & graham 1982). subfossil finds. the skagen area, holocene, recorded from the atlantic and subatlantic. melanella alba (da costa 1778) fig. 25 distribution. from norway off the lofoten islands and south to the mediterranean. occurrence. the boreal and lusitanian regions. habitat. from 16 to 135 m deep on muddy sand and gravel bottoms, an ectoparasite of holothurians. only subfossil finds. the skagen well area, holocene, recorded from the subboreal. hemiaclis ventrosa (jeffreys ms fricle 1874) fig. 26 distribution. from west and south iceland, norway off the lofoten islands, and south to the bay of biscay. occurrence. the boreal and lusitanian regions. habitat. from 100 to 300 m deep on soft bottoms. only subfossil finds. the skagen well area, holocene, recorded from the subatlantic. fig. 23. graphis albida (kanmacher 1798). skagen 4, 25.0–25.5 m b.s., lab. no. 350,93. × 40. mguh 25333. fig. 24. melanella lubrica (monterosato 1891). skagen 3, 83.73– 83.83 m b.s., lab. no. 512,93. × 40. mguh 25334. fig. 25. melanella alba (da costa 1778). skagen 3, 74.89–75.00 m b.s., lab. no. 509,93. × 40. mguh 25335. geus bulletin no 3.pmd 28-06-2004, 08:4543 44 pelseneeria stylifera (turton 1826) distribution. from norway off the lofoten islands and south to the mediterranean. few recorded from the skagerrak and kattegat, including øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. the animals are confined to the surface of regular sea urchins (fretter & graham 1982, p. 431). subfossil finds. none. enteroxenos oestergreni bonnevie 1902 distribution. recorded from scandinavia, a single danish record (jensen & knudsen 1995). occurrence. the boreal region. habitat. a parasite in the holothurian stichopus tremulus (jensen & knudsen 1995). subfossil finds. none. order neogastropoda nucella lapillus (linnaeus 1758) distribution. w greenland, around iceland, norway from north of lofoten, and south to the straits of gibraltar. the species reaches into the skagerrak, but it is uncommon in danish waters (fretter & graham 1984). however, it occurs on breakwaters along the north sea and skagerrak coasts. there are a few records from the southern kattegat and øresund (jensen & knudsen 1995). occurrence. the subarctic, boreal and lusitanian regions. habitat. intertidal on rocky shores and extends, albeit rarely, to depths of 30 to 40 m. it avoids very weedy shores and seems to stand only limited reduction of salinity (fretter & graham 1984, p. 445). subfossil finds. the kattegat, limfjord (exposed towards the skagerrak), north sea and vendsyssel regions, holocene. boreotrophon clathratus (linnaeus 1767) distribution. spitsbergen, e and w greenland, around iceland, the faroes and the coast of norway south to the skagerrak and kattegat. occurrence. the arctic, subarctic and boreal regions. habitat. from 8 m to over 1000 m on soft bottoms. subfossil finds. the vendsyssel area, late weichselian (the younger yoldia sea). boreotrophon truncatus (ström 1768) distribution. spitsbergen, e and w greenland, around iceland, norway south to the biscay. the species extends in to danish waters south to the øresund. occurrence. the arctic, subarctic, boreal and lusitanian regions. habitat. from the laminarian zone to depths of about 200 m on bottoms of a stony, gravelly or muddy nature. subfossil finds. none. ocenebra erinacea (linnaeus 1758) distribution. from the southern coasts of britain south to the mediterranean. however, the few records from fig. 26. hemiaclis ventrosa (jeffreys ms fricle 1874). skagen 4, 30.0–30.5 m b.s., lab. no. 355,93. × 20. mguh 25336. geus bulletin no 3.pmd 28-06-2004, 08:4544 45 danish waters might have been introduced with oysters (jensen & knudsen 1995). occurrence. the lusitanian region. habitat. sublittoral to 150 m deep on stony bottoms. subfossil finds. none. trophonopsis barvicensis (johnston 1825) distribution. w and s iceland, norway and south to the british isles, and further south (france) at greater depths. in danish waters from the kattegat, including the øresund, although rare (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. sublittoral at a few metres’ depth at the northern end of its range to 300–400 m at the southern end. subfossil finds. none. buccinum undatum linnaeus 1758 fig. 27 distribution. spitsbergen, w greenland, around iceland, norway and south to the bay of biscay (thorson 1944b). the species extends into the kattegat, limfjord, and the bælt sea with the mecklenburger bucht as the easternmost position. occurrence. the subarctic, boreal and lusitanian regions. habitat. sublittorally to about 1200 m deep usually on soft bottoms. subfossil finds. the kattegat, limfjord, north sea, vendsyssel and skagen areas, holocene, recorded as a fragment from the skagen well from the subatlantic. during the eemian in the kattegat, north sea, and vendsyssel regions. from the vendsyssel area found during the late weichselian (the younger yoldia sea). buccinum cyaneum bruguière 1792 distribution. spitsbergen, w and e greenland around iceland and norway north of lofoten. occurrence. the arctic, subarctic and northern part of the boreal regions. habitat. from 0 to 392 m on all sorts of bottoms – sand, clay, stones and algae (thorson 1944b). only subfossil finds. from the vendsyssel area recorded during the late weichselian (the younger yoldia sea). colus gracilis (da costa 1778) distribution. s and w iceland (empty shells), norway off the lofoten islands and south to portugal. the species extends into the kattegat. occurrence. the boreal and lusitanian regions. habitat. usually from 30 to 800 m deep (less common and deeper in the south). subfossil finds. none. colus jeffreysianus (fischer 1868) distribution. from norway south to the mediterranean. the species extends through the skagerrak to the kattegat, including the northern part of the øresund. occurrence. the boreal and lusitanian regions. habitat. from 30 to 2000 m deep on soft bottoms. subfossil finds. none.fig. 27. buccinum undatum linnaeus 1758. geus collection. limfjord, denmark. height 70 mm. mguh 25337. geus bulletin no 3.pmd 28-06-2004, 08:4545 46 colus sabini (gray 1824) distribution. from w greenland (empty shells) and s and w iceland, northern north sea and extending into the skagerrak (jensen & knudsen 1995). occurrence. the boreal region. (a subarctic extension is not considered, since only empty shells have been found off w greenland and no occurrences on n and e iceland (thorson 1941, 1944b).) habitat. from 35 to 1500 m deep on muddy bottoms. subfossil finds. none. liomesus ovum (turton 1825) distribution. from greenland, the faeroes and the coasts of norway. according to fretter & graham (1984, p. 465), the species is not recorded from the skagerrak or kattegat; however, jensen & knudsen (1995) mentioned this species as occurring in danish waters. furthermore, the species liomesus ovum cannot be found in thorson (1944b), who has buccinum ovum middendorff, which is not the same according to fretter & graham (1962), and the occurrence off the faeroes cannot be confirmed in spärck & thorson (1931). occurrence. the boreal region with only uncertain outposts into the subarctic and lusitanian regions. habitat. from 70 to 400 m deep on soft bottoms. subfossil finds. none. neptunea antiqua (linnaeus 1758) distribution. from southern norway south to the bay of biscay. the species extends into the kattegat, øresund (jensen & knudsen 1995), and the bælt sea as far east as lübecker bucht. the species has also been recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from 15 to 1200 m on all kinds of bottom, mainly soft. subfossil finds. the vendsyssel area, holocene. neptunea despecta (linnaeus 1758) distribution. from spitsbergen, e and w greenland, around iceland and the coasts of norway, south to the seas off denmark according to fretter & graham (1984), but not mentioned by jensen & knudsen (1995) in their annotated check list of recent marine molluscs of danish waters. occurrence. the arctic, subarctic and boreal regions. habitat. from 6 to 1400 m on soft bottoms. only subfossil finds. the vendsyssel region during the late weichselian (the younger yoldia sea). turrisipho moebii (dunker & metzger 1874) distribution. from the coasts of northern and southern norway and the faeroes (sipho sarsi in spärck & thorson 1931). the species extends into the skagerrak. occurrence. the boreal region. habitat. from 200 m to 1000 m deep on soft bottoms. subfossil finds. none. hinia incrassata (ström 1768) distribution. s and w iceland, norway north of lofoten, and south to the mediterranean (thorson 1941). the species extends into the kattegat and øresund, although rare in this place (jensen & knudsen 1995), and it has not been found in the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. rocky coasts in the lower part of the tidal zone. mainly found in the shallow sublittoral, but may extend to about 200 m. subfossil finds. the limfjord and vendsyssel regions, holocene. recorded from the vendsyssel region during the eemian. hinia pygmaea (lamarck 1822) fig. 28 distribution. from norway off the lofoten islands and geus bulletin no 3.pmd 28-06-2004, 08:4546 47 south to the mediterranean. it extends into danish waters such as the skagerrak, limfjord and kattegat and øresund. occurrence. the boreal and lusitanian regions. habitat. from 1 m to about 200 m on sandy bottoms. subfossil finds. the limfjord, north sea, vendsyssel and skagen areas, holocene. recorded from the subatlantic in the skagen well material. found in the bælt sea, kattegat and north sea regions during the eemian. hinia reticulata (linnaeus 1758) fig. 29 distribution. from norway off lofoten and south to the mediterranean. in danish waters within the limfjord and kattegat with fjords, the øresund and the bælt sea. occurrence. the boreal and lusitanian regions. habitat. from the tidal zone to about 15 m deep on soft bottoms. subfossil finds. the bælt sea, baltic, kattegat, limfjord, north sea, vendsyssel and skagen regions, holocene. in the skagen well material recorded from the subatlantic. from the bælt sea, baltic, kattegat, north sea and vendsyssel regions recorded from the eemian. troschelia bernicensis (king 1846) distribution. from norway north of lofoten and south to the west coast of scotland and the dogger bank. recorded fromdanishwaters by jensen&knudsen (1995). occurrence. the boreal and lusitanian (northern part) regions. habitat. lives on the continental shelves and upper slopes, between 90 and 2700 m (poppe & goto 1991). subfossil finds. none. cytharella coarctata (forbes 1840) distribution. from norway off the lofoten islands (thorson 1941: mangelia costata) and south to the mediterranean. the species is recorded from the skagerrak and kattegat where it extends into the øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone to 250 m deep on sandy bottoms. subfossil finds. the vendsyssel area, holocene. during the eemian in the north sea region. oenopota incisula (verrill 1882) distribution. the species is known from the boreal zone of the east coast of north america, and is comfig. 28. hinia pygmaea (lamarck 1822). skagen 4, 12.0–12.5 m b.s., lab. no. 337,93. × 9.6. mguh 25338. fig. 29. hinia reticulata (linnaeus 1758). skagen 3, 55.10–55.30 m b.s., lab. no. 716,93. × 4.8. mguh 25339. geus bulletin no 3.pmd 28-06-2004, 08:4547 48 mon off west greenland (posselt & jensen 1898), but it is not recorded from iceland or the coasts of the north-east atlantic (thorson 1941). occurrence. the arctic, subarctic, and in north america into the boreal regions. habitat. in the canadian north-east region the species has been collected from 6–7 to 140 m deep on clay (macpherson 1971). only subfossil finds. from the vendsyssel region recorded during the eemian and early/middle weichselian (the older yoldia clay). oenopota trevelliana (turton 1834) distribution. spitsbergen, e and w greenland, around iceland (thorson 1941: bela trevelliana (turton)), norway from north of lofoten, and south to the british isles. the species extends south to the kattegat, including øresund. occurrence. the arctic, subarctic, boreal and lusitanian (northern part) regions. habitat. sublittorally from 25 m to depths over 300 m on fine sand. subfossil finds. the vendsyssel area during the eemian. oenopota turricola (montagu 1803) fig. 30 distribution. greenland, iceland, faeroes, norway south to scotland (spärck & thorson 1931). the species extends into the limfjord and kattegat (fretter & graham 1984) and the øresund (jensen & knudsen 1995). occurrence. the arctic, subarctic and boreal regions. habitat. sublittoral from 20 to 200 m on sandy bottoms. subfossil finds. the limfjord, vendsyssel and skagen areas, holocene. from the skagen well recorded from the subatlantic. found in the vendsyssel area from the late weichselian (the younger yoldia sea). oenopota violacea (mighels & adams 1842) distribution. e and w greenland, spitsbergen, around iceland and south along the coast of norway, but not reaching the british isles (thorson 1941). occurrence. the arctic, subarctic and boreal regions. habitat. from 1 m to 761 m on mud and stones. only subfossil finds. the vendsyssel area during the eemian. bela exarata g.o. sars 1878 distribution. e and w greenland, spitsbergen, around iceland, the faeroes, and norway from north of lofoten and west of ireland. the species is also recorded west of iceland, where it has been found at depths down to 2214 m (thorson 1941). occurrence. the arctic, subarctic, boreal regions with lusitanian outposts. habitat. in the northern part of its range it belongs to the shallow-water species (norway and e greenland from 3 m) (thorson 1941). only subfossil finds. from the eemian in the vendsyssel area. mangelia attenuata (montagu 1803) distribution. from norway off the lofoten islands and fig. 30. oenopota turricola (montagu 1803). skagen 4, 21.0– 21.5 m b.s., lab. no. 346,93. × 9.6. mguh 25340. geus bulletin no 3.pmd 28-06-2004, 08:4548 49 south to the mediterranean. the species extends into the skagerrak and kattegat, including the øresund. occurrence. the boreal and lusitanian regions. habitat. from 5 to 150 m deep on sand or clay bottoms. subfossil finds. none. mangelia brachystoma (philippi 1844) fig. 31 distribution. from norway off the lofoten islands and south to the mediterranean. the species extends into the skagerrak and kattegat, including the øresund. occurrence. the boreal and lusitanian regions. habitat. sublittorally from 4 m to 60 m deep on bottoms of sand and sandy mud. subfossil finds. the skagen area, holocene, recorded from the subatlantic in the skagen well cores. during the eemian found in the vendsyssel region. mangelia nebula (montagu 1803) distribution. norway off the lofoten islands and south to the mediterranean. the species is recorded a few times from the kattegat and extends into the øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from 10 to 50 m deep on sandy bottoms. subfossil finds. none. raphitoma purpurea (montagu 1803) distribution. from northern norway off the lofoten islands south into the mediterranean. occurrence. the boreal and lusitanian regions. habitat. from 10 to 100 m deep on sandy, gravelly and stony bottoms. only subfossil finds. the limfjord region, holocene. raphitoma asperrima (brown 1827) distribution. from the coast of norway south to the mediterranean, extending into the kattegat, including the øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from 20 to 100 m deep on sandy bottoms. subfossil finds. none. raphitoma leufroyi (michaud 1821) distribution. from the coast of norway off the lofoten islands and south to the mediterranean, extending into the skagerrak and with a few records from the northern kattegat (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone to 150 m deep on sandy, shelly and stony bottoms. subfossil finds. none. raphitoma linearis (montagu 1803) distribution. sw and nw iceland, norway from north of lofoten and south to the mediterranean. the species extends into the skagerrak and kattegat and øresund, although rare (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. fig. 31. mangelia brachystoma (philippi 1844). skagen 3, 58.04– 58.25 m b.s., lab. no. 717,93. × 9.6. mguh 25341. geus bulletin no 3.pmd 28-06-2004, 08:4549 50 habitat. from 10 to 200 m deep on sandy, shelly and stony bottoms. subfossil finds. the limfjord and vendsyssel area, holocene. during the eemian found in the vendsyssel area (the turritella terebra zone in the skærumhede sequence). taranis borealis bouchet & warén 1980 distribution. so far as known, this species is confined to waters off western norway and the skagerrak (fretter & graham 1984, p. 548). occurrence. the boreal region. habitat. from 150 m to nearly 2000 m deep on soft bottoms. subfossil finds. none. taranis moerchi (malm 1861) distribution. from norway north of lofoten south to the mediterranean, extending into the kattegat. occurrence. the boreal and lusitanian regions. habitat. from 80 m deep near the northern limits of its range to over 2000 m elsewhere on soft bottoms. subfossil finds. none. admete viridula (fabricius 1780) distribution. spitsbergen, e and w greenland, around iceland, empty shells from off the faeroes, norway from north of the lofoten islands, and south to the northern borders of the north sea. included in the recent danish fauna (jensen & knudsen 1995). occurrence. the arctic, subarctic and boreal regions. habitat. from a few metres to depths of 1000 m, the greatest depths in the south of its range (fretter & graham 1984, p. 507) on soft bottoms. subfossil finds. recorded from the early/middle weichselian (the portlandia arctica zone in the skærumhede sequence) in the vendsyssel region. subclass heterobranchia order heterostropha omalogyra atomus (philippi 1841) distribution. w greenland, around iceland and norway, south to the mediterranean. known from only a few places in danish waters. occurrence. the subarctic, boreal and lusitanian regions. habitat. from the lower part of the shore to a depth of 20 m, occurring on seaweeds. only subfossil finds. the limfjord region, holocene. brachystomia carozzai van aartsen 1987 distribution. from the southern part of norway (spärck & thorson 1931) and south to the mediterranean. the species extends into the kattegat and limfjord. as commented on by fretter et al. (1986, p. 605) and jensen & knudsen (1995), the determination of these small snails living ectoparasitically on other marine organisms is still in progress, so the actual situation for the record of subfossil material should be taken with great precaution. occurrence. the boreal and lusitanian regions. habitat. from the tidal zone, where it occurs in crevices, to depths of about 70 m. subfossil finds. none. brachystomia eulimoides hanley 1844 distribution. from norway off the lofoten islands and south to the mediterranean. recorded from the limfjord, according to jensen & knudsen (1995) the only danish record, following fretter et al. (1986, p. 602). occurrence. the boreal and lusitanian regions. habitat. most frequently found on living animals of pecten, chlamys, oysters and turritella to depths of 120 m. subfossil finds. the kattegat, limfjord, north sea and geus bulletin no 3.pmd 28-06-2004, 08:4550 51 vendsyssel regions, holocene. from the north sea region during the eemian. odostomia scalaris macgillivray 1843 distribution. from southern norway south to the mediterranean. the species extends into the limfjord, through the kattegat, including the øresund, and the bælt sea into the kiel bay (fretter et al. 1986, p. 600). occurrence. the boreal and lusitanian regions. habitat. associated primarily with banks of mytilus edulis, but also recorded from other hosts (fretter et al. 1986, p. 600). subfossil finds. the limfjord and vendsyssel regions, holocene. from the bælt sea and north sea during the eemian. chrysallida decussata (montagu 1803) fig. 32 distribution. mainly southern distribution, with the shetlands as the northernmost post, but recorded from the øresund (jensen & knudsen 1995) and there are older records from east of scotland (fretter et al. 1986). occurrence. the boreal and lusitanian regions. habitat. from 14 to 40 m deep on sandy and shelly bottoms. subfossil finds.thelimfjordandskagenareas,holocene, in the skagen well recorded from the subatlantic. chrysallida eximia (jeffreys 1849) distribution. sw and nw iceland, norway north of lofoten, and south to western scotland. there is no record from danish waters. occurrence. the boreal and lusitanian (northernmost) region. habitat. from 20 m to more than 1000 m, the greater depths in the southern part of its range, on soft gravelly bottoms. only subfossil finds. the limfjord region, holocene. from the vendsyssel area during the eemian. chrysallida indistincta (montagu 1808) distribution. from norway off the lofoten islands (thorson 1941) and south to the mediterranean. it extends into the kattegat and øresund (jensen & knudsen 1995), but is not recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from 7 to 100 m deep on sandy bottoms. subfossil finds. the limfjord, north sea and vendsyssel areas, holocene. recorded from the north sea during the eemian. chrysallida obtusa (brown 1827) distribution. from norway off the lofoten islands and south to the mediterranean. the species extends into the limfjord and through the kattegat into the øresund (fretter et al. 1986, p. 562). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone in rock pools to 90 m deep in stony places, associated with oysters (fretter et al. 1986, p. 562). subfossil finds. the limfjord and vendsyssel regions, holocene. from the bælt sea and north sea regions during the eemian. fig. 32. chrysallida decussata (montagu 1803). skagen 4, 15.0– 15.5 m b.s., lab. no. 340,93. × 20. mguh 25342. geus bulletin no 3.pmd 28-06-2004, 08:4551 52 chrysallida spiralis (montagu 1803) distribution. from norway north of lofoten and south to the mediterranean. according to fretter et al. (1986, p. 574), the species extends into the kattegat and øresund, but it is absent from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone to about 120 m deep, often abundant in the neighbourhood of tubes of sedentary polychaetes (fretter et al. 1986). subfossil finds. the kattegat, limfjord, north sea and vendsyssel regions, holocene. the bælt sea, kattegat, and north sea regions during the eemian. ebala nitidissima (montagu 1803) distribution. from south of norway to the mediterranean. recorded from the kattegat region with fjords and the øresund and the bælt sea regions as far as kiel bay (fretter et al. 1986, p. 630), but absent from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from 5 to 50 m deep on muddy sand or shelly bottoms. subfossil finds. the limfjord and vendsyssel regions, holocene. from the bælt sea and north sea areas during the eemian. eulimella laevis (brown 1827) distribution. from norway off the lofoten islands and south to the mediterranean. it has not been recorded from danish waters except the øresund region (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from 20 to 400 m deep on muddy sand. subfossil finds. the limfjord, north sea and vendsyssel regions, holocene. eulimella scillae (scacchi 1835) fig. 33 distribution. from norway off the lofoten islands, but empty shells only off sw iceland (thorson 1941), and south to the mediterranean. it extends into the kattegat along the swedish west coast (fretter et al. 1986), but occurs in the øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from 20 to 400 m deep on muddy sand or sand. subfossil finds. the limfjord and skagen areas, holocene, recorded in the skagen well from the atlantic and subboreal. from the vendsyssel area found during the eemian. ondina divisa (j. adams 1797) distribution. from s and w iceland, northern norway (w finmarken) (thorson 1941) and south to the biscay. the species extends through the kattegat to the øresund (fretter et al. 1986, p. 582). occurrence. the boreal and lusitanian regions. habitat. from 18 to 200 m deep on sandy and gravelly mud. subfossil finds. the limfjord region, holocene. from the vendsyssel area found during the eemian. fig. 33. eulimella scillae (scacchi 1835). skagen 3, 82.34–82.50 m b.s., lab. no. 729,93. × 9.6. mguh 25343. geus bulletin no 3.pmd 28-06-2004, 08:4552 53 ondina obliqua (alder 1884) distribution. according to fretter et al. (1986, p. 586) from southern scandinavia to biscay. however, spärck & thorson (1931) only mentioned scotland. furthermore, only western localities have been reported from the british isles (fretter et al. 1986), and it is doubtful whether the danish records actually refer to this species. therefore, the species is here considered to be lusitanian and should not be taken as present in the recent danish fauna, although recorded by jensen & knudsen (1995). habitat. from 30 to 60 m deep in gravelly or sandy mud. subfossil finds. none. ondina diaphana (jeffreys 1848) distribution. according to spärck & thorson (1931), present from southern norway and south to the mediterranean. in danish waters the kattegat and the øresund. occurrence. the boreal and lusitanian regions. habitat. from 20 to 90 m deep on soft bottoms. subfossil finds. the limfjord and north sea regions, holocene. liostomia clavula (lovén 1846) distribution. from southern norway south to the mediterranean. extends through the kattegat to the øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from 30 to 90 m deep on soft bottoms associated with pennatula (fretter et al. 1986, p. 590). subfossil finds. none. liostomia afzelii warén 1991 this species, newly established, will not be considered further. odostomia acuta jeffreys 1848 distribution. from norway north of lofoten and south to the mediterranean. the species is found in the skagerrak, kattegat and øresund (fretter et al. 1986, p. 613). occurrence. the boreal and lusitanian regions. habitat. from 20 to 30 m deep. perhaps associated with bryozoans. subfossil finds. the limfjord region, holocene. odostomia conoidea winckworth 1932 fig. 34 distribution. from norway off the lofoten islands and south to the mediterranean. recorded from the skagerrak (fretter et al. 1986) and the øresund (jensen & knudsen 1995), but not from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from 10 to 150 m deep, usually in association with the starfish astropecten irregularis (fretter et al. 1986, p. 617). subfossil finds. the bælt sea, limfjord, north sea, vendsyssel and skagen regions, holocene. in the skagen area recorded from the subatlantic. fig. 34. odostomia conoidea winckworth 1932. skagen 3, 48.90– 49.10 m b.s., lab. no. 714,93. × 40. mguh 25344. geus bulletin no 3.pmd 28-06-2004, 08:4553 54 odostomia turrita hanley 1844 distribution. from norway north of lofoten and south to the mediterranean. the species extends through the kattegat to the øresund (fretter et al. 1986, p. 612), but is not recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone to 100 m deep on weed and clay bottoms. subfossil finds. the limfjord and vendsyssel regions, holocene. recorded from the vendsyssel area during the eemian. odostomia albella lovén 1846 distribution. empty shells recorded from spitsbergen, sw and nw iceland, norway north of lofoten and south to the mediterranean. the species occurs in the skagerrak (fretter et al. 1986), the øresund area (jensen & knudsen 1995), and the limfjord region (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone to depths of about 100 m on boulders, associated with growth of pomatoceros (fretter et al. 1986). subfossil finds. the limfjord, north sea and vendsyssel regions, holocene. from the north sea region during the eemian. odostomia plicata (montagu 1803) distribution. from the southern part of scandinavia (not southern norway (spärck & thorson 1931)) and south to the mediterranean (fretter et al. 1986, p. 610). the species extends into the kattegat, including the øresund (jensen & knudsen 1995), but is not recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. in the tidal zone associated with pomatoceros triqueter (fretter et al. 1986). subfossil finds. the limfjord and vendsyssel regions, holocene. odostomia umbilicaris (malm 1863) fig. 35 distribution. from southern norway (spärck & thorson 1931) to the british isles. the species extends to the swedish west coast but not further into the kattegat (fretter et al. 1986, p. 620). occurrence. the boreal and lusitanian (northern part) regions. habitat. from 20 to 275 m deep, found on the bivalve mytilus adriaticus (fretter et al. 1986). subfossil finds. the skagen area, holocene, recorded from the atlantic and subatlantic. turbonilla crenata (brown 1827) distribution. from norway off the lofoten islands (thorson 1941) and south to the mediterranean. the species extends into the kattegat and øresund (fretter et al. 1986), but is absent from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from 15 to 350 m deep on fine sand. subfossil finds. the limfjord and north sea regions, holocene. recorded from the bælt sea and north sea during the eemian. fig. 35. odostomia umbilicaris (malm 1863). skagen 3, 82.34–82.50 m b.s., lab. no. 725,93. × 20. mguh 25345. geus bulletin no 3.pmd 28-06-2004, 08:4554 55 turbonilla delicata monterosato 1884 fig. 36 distribution. from the south-western part of the british isles to the mediterranean (fretter et al. 1986, p. 636). however, recorded from the northern kattegat and northern øresund although, rare (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. on soft bottoms, but at uncertain depths (fretter et al. 1986, p. 636). subfossil finds. the limfjord, north sea and skagen regions, holocene. from the skagen area recorded from the subatlantic. turbonilla lactea (linnaeus 1758) distribution. from northern norway and south to the mediterranean. present although uncommon in danish waters (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone to depths of about 80 m, occurring under stones in silty places in the tidal zone and on soft, muddy and sandy bottoms sublittorally (fretter et al. 1986, p. 634). subfossil finds. the limfjord, north sea and vendsyssel regions, holocene. recorded from the bælt sea, kattegat, and north sea regions during the eemian. subclass opisthobranchia order bullomorpha acteon tornatilis (linnaeus 1758) distribution. w and s iceland, norway off the lofoten islands and south to the mediterranean (lemche 1938). the species extends into the kattegat, including the øresund (petersen 1888, p. 78), but it is not recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone down to 250 m in sand, usually in sheltered areas (poppe & goto 1991, p. 192). subfossil finds. the limfjord, north sea and vendsyssel regions, holocene. recorded from the north sea and vendsyssel areas during the eemian. haminoea navicula (da costa 1778) distribution. from the british isles south to the mediterranean. occurrence. the lusitanian region. habitat. lives in the zostera beds in sheltered areas (poppe & goto 1991, p. 196). only subfossil finds. the bælt sea, kattegat, and north sea regions during the eemian. cylichna cylindracea (pennant 1777) distribution. w and s iceland, norway off the lofoten islands and south to the mediterranean. the species extends into the limfjord and kattegat, including øresund (petersen 1888, p. 78; jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from 40 to 200 m deep, in sand. according to lemche (1938) associated with the boreal amphiura filiformis community in deeper water south of iceland, corresponding to its occurrence in danish waters. subfossil finds. the limfjord and vendsyssel regions, holocene. fig. 36. turbonilla delicata (monterosato 1874). skagen 4, 20.0– 20.5 m b.s., lab. no. 345,93. × 20. mguh 25346. geus bulletin no 3.pmd 28-06-2004, 08:4555 56 cylichna alba (brown 1827) fig. 37 distribution.wand e greenland, around iceland, norway north of lofoten and south to the bay of biscay at greater depth (lemche 1938, p. 9). the species extends into the north sea and skagerrak (lemche 1928, p. 4). occurrence. the arctic, subarctic, boreal and lusitanian regions. habitat. according to lemche (1928), this species is also widely distributed in depth, being found in low water to depths down to 2700 m, on clay bottom (faeroes). subfossil finds. the limfjord, north sea and skagen areas, holocene. in the skagen well recorded from the subatlantic. also found in the vendsyssel area from the early/middle and late weichselian (the older and younger yoldia sea respectively). cylichna occulta (mighels 1841) distribution. e and w greenland, n and e iceland, norway north of lofoten (lemche 1938). occurrence. the arctic, subarctic and boreal (northernmost part) regions. will be referred as an arctic and subarctic species. habitat. from 10 to 388 m deep (iceland), especially found within the macoma calcarea community, and might also occur within the yoldia hyperborea community (lemche 1938). only subfossil finds. recorded from the kattegat region during the early/middle weichselian and the vendsyssel region during the early/middle and late weichselian (the older and younger yoldia sea respectively). scaphander lignarius (linnaeus 1758) distribution. from s and w iceland, the faeroes, norway off the lofoten islands and south to the mediterranean. occurrence. the boreal and lusitanian regions. habitat. from 60 to 700 m deep. the species is associated with the boreal spisula elliptica community on sandy plateaus south and west of iceland (lemche 1938, p. 7). subfossil finds. none. scaphander punctostriatus (mighels & adams 1841) distribution. e and w greenland, s and w iceland, the faeroes, norway north of lofoten and south to the mediterranean. occurrence. the arctic, subarctic, boreal and lusitanian regions. habitat. recorded from depths between 10 and 3000 m, this probably having some relation to its wide horizontal distribution. around the faeroes the species may be expected to be found on the great, sandy plateaus (lemche 1928, p. 4). subfossil finds. none. philine aperta (linnaeus 1767) distribution. from the faeroes, western norway off lofoten, and south to the mediterranean (lemche 1928). the species extends through kattegat and øresund, into the bælt sea as far as kieler bugt (petersen 1888, p. 83), and it is also recorded from the limfjord (petersen 1986a). fig. 37. cylichna alba (brown 1827). skagen 3, 44.88–45.00 m b.s., lab. no. 499,93. × 9.6. mguh 25347. geus bulletin no 3.pmd 28-06-2004, 08:4556 57 occurrence. the boreal and lusitanian regions. habitat. the species prefers shallow water but has been found at depths down to 100 m off the faeroes (lemche 1928) on sandy bottoms. in danish waters recorded from 10 to 30 m, also on sandy bottoms (petersen 1888). subfossil finds. the limfjord and vendsyssel regions, holocene. recorded from the bælt sea and the north sea during the eemian. philine angulata jeffreys 1867 this species is mentioned by jensen & knudsen (1995), but only on the basis of one broken shell found in the zoological museum in copenhagen; difficult to identify; occurrence uncertain. so in the light of “the difficult problem of the relation between punctata and angulata”, as treated by lemche (1948, p. 67), this find will be omitted, also that no subfossil species have been found. philine catena (montagu 1803) fig. 38 distribution. from norway off the lofoten islands and south to the mediterranean. occurrence. the boreal and lusitanian regions. habitat. from low water-mark to 76 m (forbes & hanley 1853) subfossil finds. the skagen region during the eemian. philine denticulata (adams 1800) distribution. from norway south to the mediterranean. occurrence. the boreal and lusitanian regions. subfossil finds. none. philine punctata (adams 1800) distribution. the faeroes, southern part of the west coast of norway, and south to the mediterranean. lemche (1928) also mentioned occurrences from greenland, which, however, was not repeated in later papers (lemche 1941a, b). jensen & knudsen (1995) report occurrences from the øresund, although rare, and petersen (1888) has a single find from the bælt sea and petersen (1986a) from the limfjord. occurrence. the boreal and lusitanian regions. habitat. at the faeroes the species is recorded from depths down to 240 m (lemche 1928). subfossil finds. the limfjord and vendsyssel regions, holocene. philine quadrata (wood 1839) distribution. w greenland, around iceland, norway north of lofoten, and south to the mediterranean. recorded from the øresund, but rare (jensen & knudsen 1995). occurrence. the subarctic, boreal and lusitanian regions. habitat. the vertical range of the species is about 35 m down to 2150 m (lemche 1938). subfossil finds. none. philine scabra (müller 1776) distribution. s and w iceland, the faeroes, norway north of lofoten, and south to the mediterranean. the species extends into the kattegat, including the øresund (petersen 1888, p. 84). occurrence. the boreal and lusitanian regions.fig. 38. philine cf. catena (montagu 1803). skagen 3, 180.77– 180.89 m b.s., lab. no. 798,93. × 40. mguh 25348. geus bulletin no 3.pmd 28-06-2004, 08:4557 58 habitat. off the coasts of w and s iceland from 20– 216 m on sandy bottom (lemche 1938, p. 11); however, in danish waters (kattegat) from 20–40 m on mixed bottom (petersen 1888, p. 84). subfossil finds. none. philinoglossa helgolandica hertling 1932 distribution. from the north sea – helgoland – and south to the mediterranean, the species might occur in danish waters (jensen & knudsen 1995, p. 29). occurrence. the boreal and lusitanian regions. habitat. probably in shell gravel (jensen & knudsen 1995). subfossil finds. none. order anaspidea diaphana minuta brown 1827 distribution. e and w greenland, around iceland, the faeroes, norway north of lofoten, and south to the mediterranean. the species extends into the kattegat and øresund (petersen 1888; jensen & knudsen 1995). occurrence. the arctic, subarctic, boreal and lusitanian regions. habitat. in general from the tidal zone to 770 m, but is said to prefer sandy clay at depths of 20–40 m (lemche 1928). this is very much the same as for the kattegat, where petersen (1888) says that the species prefers mixed bottom at a depth of about 19–38 m. subfossil finds. the limfjord and vendsyssel regions, holocene. retusa obtusa (montagu 1803) distribution. e and w greenland, around iceland, the faeroes, norway north of lofoten, and south to the british isles (shetland and scotland) (lemche 1928). the species is found in the limfjord and is common in the fjords and bays bordering the kattegat, including the øresund, and extends into the bælt sea and the western part of the baltic (petersen 1888, p. 81). occurrence. the arctic, subarctic and boreal regions. habitat. from the intertidal zone down to 300 m deep in mud or fine sand. in danish waters petersen (1888) points to the observed differences in depth, i.e. in the kattegat region around 20 m while 60 m deep in the baltic. the species is connected with the arctic macoma community (lemche 1941a, b). subfossil finds. the bælt sea, kattegat, limfjord, north sea and vendsyssel regions, holocene. also recorded from the vendsyssel area during the early/middle weichselian and the late weichselian (the older yoldia sea and the younger yoldia sea respectively). retusa truncatula (bruguière 1792) fig. 39 distribution. the faeroes, norway north of lofoten, and south to the mediterranean (lemche 1928, 1938). the species extends through the kattegat, including the øresund, and into the bælt sea, and it is also recorded from the limfjord (petersen 1888, p. 80). occurrence. the boreal and lusitanian regions. habitat. in general the species lives from the tidal zone down to 200 m (poppe & goto 1991). however, in danish waters, according to petersen (1888), it lives in shallow water down to only about 20 m deep on sandy bottoms with zostera. subfossil finds. the bælt sea, baltic, kattegat, limfjord, north sea, vendsyssel and skagen regions, holocene. fig. 39. retusa truncatula (bruguière 1792). skagen 3, 67.34–67.39 m b.s., core sample k-25. × 9.6. mguh 25349. geus bulletin no 3.pmd 28-06-2004, 08:4558 59 recorded from the skagen area from the subboreal and subatlantic. from the bælt sea and the north sea recorded during the eemian. retusa umbilicata (montagu 1803) fig. 40 distribution. from norway north of lofoten (lemche 1928) and south to the mediterranean (poppe & goto 1991). the species extends into the kattegat and øresund (petersen 1888; jensen & knudsen 1995), but is not recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from depths of about 20 to 30 m on mixed bottoms (petersen 1888). subfossil finds. the limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen well recorded from the subboreal and subatlantic. from the bælt sea,north sea, and vendsyssel regions found during the eemian. rhizorus acuminatus (bruguière 1792) distribution. from western and southern norway south to the mediterranean, in danish waters from the southern kattegat. occurrence. the boreal and lusitanian regions. habitat. from the tidal zone to 800 m deep. subfossil finds. none. akera bullata o.f. müller 1776 distribution. the faeroes, norway north of lofoten, and south to the mediterranean (lemche 1928). the species extends into the limfjord, the kattegat, including the øresund, and the bælt sea as far as the kieler bugt (petersen 1888). occurrence. the boreal and lusitanian regions. habitat. in sheltered bays down to 370 m deep (poppe & goto 1991). common in shallow water with zostera (petersen 1888). subfossil finds. the bælt sea, kattegat, limfjord and vendsyssel regions, holocene. in the bælt sea and kattegat regions recorded from the eemian. order thecosomata limacina retroversa (fleming 1823) fig. 41 distribution.eand w greenland, around iceland, norway north of lofoten, and south to ireland (lemche 1938). it is common in the north sea and skagerrak, penetrating into the kattegat, occasionally even into the bælt sea (kramp 1961). occurrence. the arctic, subarctic, boreal and lusitanian (northernmost) regions. habitat. pelagic. subfossil finds. recorded from the vendsyssel and skagen regions during the eemian and from the vendsyssel region during the late weichselian (the younger yoldia sea). fig. 40. retusa umbilicata (montagu 1803). skagen 3, 70.10–70.30 m b.s., lab. no. 721,95. × 20. mguh 25350. fig. 41. limacina retroversa (fleming 1823). skagen 3, 180.57– 180.70 m b.s., lab. no. 797,93. × 20. mguh 25351. geus bulletin no 3.pmd 28-06-2004, 08:4559 60 order gymnosomata clione limacina (phipps 1774) fig. 42 distribution. e and w greenland, around iceland, the faeroes, norway north of lofoten (lemche 1938), and south to the west coast of england. common in the northern part of the north sea and in the skagerrak, occasionally penetrating into the kattegat (kramp 1961). rare occurrences in the øresund (jensen & knudsen 1995). occurrence. the arctic, subarctic, boreal and lusitanian (northern part) regions. habitat. pelagic. subfossil finds. recorded (imprint) from the skagen well during the late weichselian (the younger yoldia sea). subclass pulmonata order basommatophora ovatella myosotis (draparnaud 1801) distribution. from scandinavia along the coast of western europe, south to the mediterranean (steenberg 1911, p. 204). in danish waters recorded from the bælt sea and the baltic (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. lives on sea wrack along the coasts (steenberg 1911). subfossil finds. none. lymnaea (radix) peregra (müller 1774) f. baltica linné distribution. lymnaea peregra is foundall overeurope: iceland, the faeroes, norway north of lofoten, and south to the mediterranean (mandahl-barth 1938). the species occurs in the baltic and southern øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. lymnaea peregra f. baltica linné and l. p. f. succinea nilsson are brackish water forms tolerating up to 8‰ salt (mandahl-barth 1949, p. 74). subfossil finds. the bælt sea and baltic regions, holocene. class scaphopoda order siphonodentalioida cadulus subfusiforme (m. sars 1865) fig. 43 distribution. western iceland, the faeroes, norway north of lofoten, and south to the mediterranean (knudsen 1949b). in danish waters taken in the skagerrak (muus 1959). occurrence. the boreal and lusitanian regions. habitat. from 80 to 1300 m deep on mud bottoms. in danish waters at a depth of 230 m on clay bottom (muus 1959). subfossil finds. recorded from the skagen well during the eemian. fig. 42. clione limacina (phipps 1774). skagen 3, 115.07–115.15 m b.s., lab. no. 749,93. × 20. mguh 25352. fig. 43. cadulus subfusiforme (m. sars 1865). skagen 3, 180.77– 180.89 m b.s., lab. no. 798,93. × 20. mguh 25353. geus bulletin no 3.pmd 28-06-2004, 08:4560 61 siphonodentalium lobatum (sowerby 1860) fig. 44 distribution. e and w greenland, n and e iceland (knudsen 1949b), the faeroes, norway north of lofoten, and south to portugal. the species might be found in norske rende. in norway often found in glacial deposits (muus 1959). occurrence. the main areas are the arctic, subarctic and boreal regions, but the species extends into the lusitanian region. habitat. from 36 to 3116 m on mud (knudsen 1949b). only subfossil finds. recorded from the skagen well in the late weichselian (the younger yoldia sea) and from the eemian. entalina tetragona (brocchi 1814) fig. 45 distribution. from northern norway south to the bay of biscay and the mediterranean. from danish waters in the north sea and skagerrak (muus 1959). occurrence. the boreal and lusitanian regions. habitat. it is a deep-water species preferring mud bottoms (poppe & goto 1993). in the danish waters known from 100 to 480 m on mixed bottom (muus 1959) connected with the amphilipis norvegica/pecten vitreus community at depths of 250–700 m. subfossil finds. recorded from the skagen well during the eemian. order dentalioida antalis agile g.o. sars 1878 distribution. s and w iceland, norway off the lofoten islands, and south to the mediterranean. in the danish waters the species extends from the north sea into the skagerrak and kattegat. occurrence. the boreal and lusitanian regions. habitat. from 55 m down to 1250 m. in scandinavian waters rarely at depths less than 70 m (muus 1959). subfossil finds. none. antalis entalis (linnaeus 1758) distribution. w greenland, iceland, the faeroes, norway north of lofoten, and south to the mediterranean (knudsen 1949b). common in the danish waters, extending into the kattegat (muus 1959) although rare in the øresund (jensen & knudsen 1995). occurrence. the subarctic, boreal and lusitanian regions. habitat. from 20 to 400 m in the kattegat region, and in the north sea between 30 and 200 m (muus 1959). however, the vertical range in general goes from 6– 3200 m (knudsen 1949b). subfossil finds. the skagen well during the eemian. fig. 44. siphonodentalium lobatum (sowerby 1860). skagen 3, 114.64–? m b.s., core sample k-57. × 9.6. mguh 25354. fig. 45. entalina tetragona (brocchi 1814). skagen 3, 183.77– 184.00 m b.s., lab. no. 799,93. × 9.6. mguh 25355. geus bulletin no 3.pmd 28-06-2004, 08:4561 62 dentalium vulgare da costa 1778 fig. 46 distribution. from the british isles and south to the mediterranean. might be found in the southern north sea. occurrence. mainly the lusitanian region. habitat. sublittorally from 1 m and down to 50 m on mud and sandy bottoms. only subfossil finds. recorded from the skagen well during the eemian. class bivalvia subclass palaeotaxodonta order nuculoida nucula nitidosa winckworth 1930 fig. 47 distribution. from norway off the lofoten islands and south to the mediterranean (madsen 1949). the species extends from the more shallow part of the north sea and skagerrak into the kattegat, including the øresund. very common in the western part of the limfjord (jensen & spärck 1934, p. 23). occurrence. the boreal and lusitanian regions. habitat. in general from 7 to 250 m deep on fine sand or sand/mud bottoms (poppe & goto 1993). however, according to jensen & spärck (1934) the species is often found at depths of 6–10 m in the limfjord and 10–30 m in the kattegat region. subfossil finds. the limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen well recorded from the subatlantic. during the eemian found in the bælt sea, baltic and north sea regions. nucula nucleus (linnaeus 1767) fig. 48 distribution. from norway off the lofoten islands, the faeroes, and south to the mediterranean (madsen 1949). posselt & jensen (1898) have but few records from west greenland. the species extends from the north sea into the kattegat, including the øresund (jensen & spärck 1934), and is recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. lives from the tidal zone down to 150 m on gravel and mud bottoms. in danish waters common between 20 and 100 m. subfossil finds. the limfjord, north sea, vendsyssel fig. 46. dentalium vulgare da costa 1778. skagen 3, 181.17– 181.21 m b.s., core sample k-99. × 9.6. mguh 25356. fig. 47. nucula nitidosa winckworth 1930. skagen 4, 21.0–21.5 m b.s., lab. no. 346,93. × 9.6. right valve. mguh 25357. fig. 48. nucula nucleus (linnaeus 1767). skagen 4, 9.0–9.5 m b.s., lab. no. 334,93. × 4.8. right valve. mguh 25358. geus bulletin no 3.pmd 28-06-2004, 08:4562 63 and skagen regions, holocene. recorded in the skagen well from the subatlantic. in the eemian recorded from the baltic and vendsyssel regions. nucula sulcata (bronn 1831) distribution. norway off the lofoten islands and south to the mediterranean. in danish waters common in the deeper parts of the kattegat and øresund (jensen & knudsen 1995), and it is also recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from 10 to 400 m deep, on mud or clay bottoms, and down to 2250 m (poppe & goto 1993). in kattegat from 50 to 100 m deep on silty bottom (jensen & spärck 1934). subfossil finds. the north sea, holocene. from the bælt sea, north sea and vendsyssel regions recorded from the eemian. nuculoma hanleyi winckworth 1931 distribution. from the british isles south to spain (poppe & goto 1993). recorded from the northern kattegat (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from 30 to 90 m deep on mud and gravel bottoms. subfossil finds. none. nuculoma tenuis (montagu 1808) distribution. w and e greenland, spitsbergen, around iceland, the faeroes, norway north of lofoten, and south to the mediterranean (madsen 1949). the species extends from the deeper part of the skagerrak into the kattegat, øresund, and bælt sea (jensen & spärck 1934). occurrence. the arctic, subarctic, boreal and lusitanian regions. habitat. from off-shore down to 300 m on muddy bottoms (poppe & goto 1993). the species is more littoral in the northern latitudes than in the south (jensen & spärck 1934). subfossil finds. from the limfjord and north sea regions, holocene. recorded from the vendsyssel area during the eemian, and from the kattegat and vendsyssel area during the early/middle weichselian (the older yoldia sea), and furthermore from the vendsyssel region during the late weichselian (the younger yoldia sea). nuculana minuta (müller 1776) fig. 49 distribution. se and w greenland, around iceland, the faeroes, norway north of lofoten, and south to the british isles. according to ockelmann (1958), lacking in the most high-arctic seas. in south-western europe only at depths greater than 400 m (madsen 1949). the species extends from the north sea and skagerrak into the kattegat and øresund (jensen & spärck 1934). occurrence. the subarctic and boreal regions. habitat. from 10 to 190 m. however, recorded from 2000 m on mud, sand and gravel bottoms (poppe & goto 1993). in danish waters common on silty bottom (jensen & spärck 1934). subfossil finds. the skagen region, holocene. recorded from the skagen well during the subboreal and subatlantic. from the vendsyssel region found during the eemian and in the skagen area during the late weichselian (the younger yoldia sea). fig. 49. nuculana minuta (müller 1776). geus collection. læsø, denmark. × 4.8. right valve. mguh 25359. geus bulletin no 3.pmd 28-06-2004, 08:4563 64 nuculana pernula (müller 1776) fig. 50 distribution. e and w greenland, around iceland, the faeroes,norwaynorthof lofoten, and south to the bay of biscay. however, in sw europe only at depths greater than 400 m (madsen 1949). in danish waters the species extends from the north sea, skagerrak, and into the kattegat, øresund, and bælt sea region north of femern (jensen & spärck 1934). occurrence. the arctic, subarctic and boreal regions. habitat. lives off-shore between 80 and 900 m deep, especially in mud bottoms (poppe & goto 1993). however, in danish waters the species is found from a depth of 20 m (kattegat) to 200 m (skagerrak) according to jensen & spärck (1934), and in the arctic (east greenland) from 3–9 m (ockelmann 1958), being mostly littoral in the arctic. subfossil finds. recorded from the vendsyssel and skagen regions during the eemian. from the kattegat, vendsyssel and skagen regions during the early and middle weichselian (the older yoldia sea). from the vendsyssel and skagen regions during the late weichselian (the younger yoldia sea). yoldia hyperborea lovén 1859 fig. 51 distribution. e and w greenland, spitsbergen, around iceland and norway north of lofoten (madsen 1949). occurrence. the arctic, subarctic and boreal (northern part – high-boreal) regions. habitat. from about 5 to 675 m on clay or mud, in few cases sand (madsen 1949). only subfossil finds. recorded from the vendsyssel and skagen areas during the early and middle weichselian (the older yoldia sea) and from the skagen well also during the late weichselian (the younger yoldia sea). portlandia arctica (gray 1824) fig. 52 distribution. according to ockelmann (1958, p. 26): “widely distributed in high-arctic seas”. occurrence. north greenland, east greenland, spitsbergen, the barents sea, novaya zemlya, the kara sea, the siberian ice sea, viz: arctic and subarctic regions. habitat. from 2 m to 340 m deep, however, the species is most common at depths between 10 and 50 m on a muddy or clayey bottom (ockelmann 1958, p. 25). fig. 50. nuculana pernula (müller 1776). geus collection. læsø, denmark. × 4.8. right valve. mguh 25360. fig. 51. yoldia hyperborea lovén 1859. geus collection. nordre strømfjord, greenland. × 4.8. right valve. mguh 25361. fig. 52. portlandia arctica (gray 1824). geus collection. disko, greenland, 46 m a.s.l. × 4.8. specimen seen from the right. mguh 25362. geus bulletin no 3.pmd 28-06-2004, 08:4564 65 only subfossil finds. recorded from the kattegat, vendsyssel, and skagen areas during the early and middle weichselian (the older yoldia sea) and from the same areas during the late weichselian (the younger yoldia sea). yoldiella lucida (lovén 1846) distribution. w greenland, around iceland, norway north of lofoten (madsen 1949), and south over the british isles to the mediterranean. in the danish waters found in the deeper part of the skagerrak (jensen & spärck 1934). occurrence. the subarctic, boreal and lusitanian regions. habitat. from about 20 m to about 1400 m (iceland), in the skagerrak from about 200 m to more than 600 m on clayey bottoms. subfossil finds. from the vendsyssel area during the eemian. yoldiella lenticula (möller 1842) fig. 53 distribution. e and w greenland, spitsbergen, siberian sea, norway north of the lofoten islands, south to the british isles. further south only at depths greater than 400 m (madsen 1949). occurrence. the arctic, subarctic and boreal regions. habitat. most common at depths between 20 and 200 m (east greenland) on rather pure clay or mud (ockelmann 1958). only subfossil finds. recorded from the vendsyssel area during the eemian, early/middle weichselian (the older yoldia sea) and late weichselian (the younger yoldia sea). found in the skagen area from the late weichselian (the younger yoldia sea). yoldiella frigida (torell 1859) fig. 54 distribution. e and w greenland, spitsbergen, around iceland, norway from north of lofoten, and south to the mediterranean (madsen 1949). in the danish waters found in the deeper parts of the skagerrak (jensen & spärck 1934). however, this is questioned by ockelmann (1958), who says that the main distribution is high-arctic and therefore the recent finds in danish waters should be referable to yoldiella nana. however, according to the old information, the species must have a wide distribution. occurrence. the arctic and subarctic regions, boreal and lusitanian. habitat. most common at depths between 30 and 150 m on bottoms consisting of clay, mud, and clay mixed with sand and gravel (ockelmann 1958). subfossil finds. following the distribution of portlandia frigida sensu jensen & spärck (1934), the species has been recorded from the vendsyssel and skagen areas during the eemian and from the early and middle weichselian (the older yoldia sea) and the late weichselian (the younger yoldia sea). fig. 53. yoldiella lenticula (möller 1842). skagen 3, 121.39– 121.50 m b.s., lab. no. 759,93. × 20. specimen seen from the left. mguh 25363. fig. 54. yoldiella frigida (torell 1859). skagen 3, 127.39–127.50 m b.s., lab. no. 763,93. × 20. specimen seen from the right. mguh 25364. geus bulletin no 3.pmd 28-06-2004, 08:4565 66 yoldiella philippiana (nyst 1845) distribution. from norway off the lofoten islands and south to the mediterranean. occurrence. the boreal and lusitanian regions. habitat. lives to a depth of about 135 m (poppe & goto 1993); however, found at depths of 2500 m (jensen & spärck 1934). only subfossil finds. recorded from the vendsyssel area during the eemian. yoldiella nana (m. sars 1846) distribution. considering the discussion by ockelmann (1958) on distinguishing between portlandia frigida and yoldiella nana, where the latter “at least in part” should be referable to p. fraterna, it is not possible to give any information on the distribution of yoldiella nana mentioned by jensen & knudsen (1995) as being part of the recent danish fauna. subfossil finds. none. malletia obtusa (g.o. sars 1872) distribution. from norway off the lofoten islands and south to sw europe. the occurrences around the british isles and in the mediterranean are at depths greater than 400 m (madsen 1949). occurrence. the boreal and lusitanian regions. habitat. this species has a wide range of depths according to madsen (op. cit.) from 20–3200 m deep. in danish waters only from the deeper part (> 300 m), and in the skagerrak mostly at a depth of 600 m (jensen & spärck 1934). subfossil finds. none. subclass pteriomorphia order arcoida acar nodulosa (müller 1766) distribution. s and w iceland, the faeroes, norway north of lofoten, and south to the mediterranean. the species has been taken on the dogger bank, but not in the inner danish waters. occurrence. the boreal and lusitanian regions. habitat. from the tidal zone down to 1000 m fixed with its byssus to hard substrates (poppe & goto 1993). subfossil finds. none. bathyarca glacialis (gray 1824) fig. 55 distribution. w and e greenland, spitsbergen, and around iceland. the occurrences from sw europe are of dead shells and from deep water (ockelmann 1958). occurrence. mainly the arctic and subarctic regions; however, the occurrences from southern iceland imply extension into the boreal region as well. habitat. from 6–10 m down to 425 m (east greenland), but most abundant below 40 m on clay bottoms with stones and gravel, where the astarte crenata community occurs. only subfossil finds. recorded from the vendsyssel (not in the skærumhede sequence) and skagen regions during the early/middleweichselian and the late weichselian respectively. bathyarca pectunculoides (scacchi 1834) distribution.wand e greenland, around iceland, norway north of lofoten, and south to the mediterranean. in the danish waters the species is rather common in the deeper part of the skagerrak between 300–700 m (jensen & spärck 1934). fig. 55. bathyarca glacialis (gray 1824). skagen 3, 114.64–? m b.s., core sample-57. × 9.6. fragments of left valve. mguh 25365. geus bulletin no 3.pmd 28-06-2004, 08:4566 67 occurrence. the arctic, subarctic, boreal and lusitanian regions. habitat. from 50 m (the shetland isles) to more than 2000 m (the mediterranean) (madsen 1949). subfossil finds. none. order mytiloida mytilaster lineatus (gmelin 1791) distribution. a mediterranean species according to jensen & spärck (1934) but also found in the neighbouring atlantic (poppe & goto 1993). occurrence. the lusitanian region. habitat. intertidal, attached to rocks. only subfossil finds. the eemian in the kattegat and north sea regions. mytilaster solidus form minimus (poli 1795) distribution. from bretagne and south into the mediterranean. occurrence. the lusitanian region. habitat. intertidal, attached to rocks or algae. subfossil finds. recorded from the bælt sea during the eemian. mytilus edulis linnaeus 1758 fig. 56 distribution. w and e greenland, but only along the south-eastern coast (ockelmann 1958), around iceland, the faeroes, norway north of lofoten, and south to the mediterranean. known from all parts of the danish waters, including the baltic. occurrence. the subarctic, boreal and lusitanian regions. habitat. intertidal to 40 m deep, but in danish waters common as an epifaunal element down to a depth of 10 m; however, in the baltic as deep as 40 m (jensen & spärck 1934). subfossil finds. the bælt sea, baltic, kattegat, limfjord, north sea, vendsyssel and skagen regions, holocene. recorded in the subatlantic from the skagen well. in the eemian records from the bælt sea, baltic, kattegat, north sea and vendsyssel regions. in the late weichselian also recorded from the vendsyssel region (the younger yoldia sea). modiolula phaseolina (philippi 1844) distribution. from sw and nw iceland, the faeroes, norway north of lofoten, and south to the british isles (petersen 1968) and the mediterranean. the species extends from the north sea and skagerrak into the kattegat and øresund (jensen & knudsen 1995), but is not recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone and down to 160 m, attached by its byssus to rocks or on the base of the larger seaweeds (poppe & goto 1993), but also recorded from depths of 1000 m. subfossil finds. the bælt sea, limfjord and vendsyssel regions, holocene. recorded from the bælt sea and the north sea regions during the eemian. modiolus adriaticus (lamarck 1819) distribution. recorded from the southern part of the british isles and denmark (petersen 1968) south to the mediterranean. in danish waters taken in the kattegat, including the øresund (jensen & spärck 1934), but fig. 56. mytilus edulis linnaeus 1758. skagen 4, 18.7–18.8 m b.s., lab. no. 311,95. × 4.8. fragment of left valve. mguh 25366. geus bulletin no 3.pmd 28-06-2004, 08:4567 68 not observed in the limfjord (petersen 1986a). it might have been passed over in many places, as mentioned by petersen (1888, p. 127). occurrence. the boreal and lusitanian regions. habitat. sublittoral between 14 and 75 m on mud bottoms. subfossil finds. the limfjord and vendsyssel region, holocene. modiolus modiolus (linnaeus 1758) distribution. around iceland, the faeroes, norway north of lofoten, and south to the british isles (petersen 1968) and the bay of biscay (poppe & goto 1993). in danish waters, including the limfjord, the species extends into the bælt sea (jensen & spärck 1934) and the øresund (jensen & knudsen 1995). occurrence. the subarctic, boreal and lusitanian regions. habitat. from the extreme low tide down to 150 m attached with its byssus to rocks or gravel (poppe & goto 1993). in danish waters the species replaces mytilus edulis as the dominating epifaunal element in deeper water (jensen & spärck 1934). subfossil finds. the bælt sea, limfjord and vendsyssel regions, holocene. recorded from the baltic and north sea during the eemian. musculus discors (linnaeus 1767) fig. 57 distribution. e and w greenland, around iceland and spitsbergen (madsen 1949), norway from north of lofoten (petersen 1968) and south to the mediterranean. in danish waters the species extends into the bælt sea and øresund (jensen & spärck 1934), but it has not been observed in the limfjord (petersen 1986a). occurrence. the arctic, subarctic, boreal, and lusitanian regions. habitat. from the intertidal zone on algae (poppe & goto 1993) and rarely on water deeper than about 200 m (jensen & spärck 1934). subfossil finds. the bælt sea, kattegat, limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen well recorded from the subatlantic. musculus laevigatus (gray 1824) distribution. e and w greenland, around iceland and norway north of lofoten. occurrence. the arctic, subarctic and boreal regions. habitat. from the infralittoral zone down to 83 m (poppe & goto 1993). only subfossil finds. recorded from the early/middle weichselian (the older yoldia sea, but not in the skærumhede sequence) and the late weichselian (the younger yoldia sea) in the vendsyssel region. musculus niger (gray 1824) distribution. e and w greenland, around iceland, the faeroes, norway north of lofoten, and south to the north sea and the irish sea (madsen 1949; petersen 1968). in danish waters the species extends from the north sea into the kattegat, øresund and bælt sea as far as warnemünde (jensen & spärck 1934; jensen & knudsen 1995). occurrence. the arctic, subarctic and boreal regions. habitat. most often found in danish waters at water depths of more than 25 m (jensen & spärck 1934). poppe & goto (1993) indicate from 7 m deep down to about 135 m. however, off the east greenland coast the species is rarely met with at depths exceeding 40 m (ockelmann 1958). fig. 57. musculus discors (linnaeus 1767). skagen 4, 28.0–28.5 m b.s., lab. no. 353,93. × 20. left valve. mguh 25367. geus bulletin no 3.pmd 28-06-2004, 08:4568 69 subfossil finds. recorded from the vendsyssel region during the eemian, the early and middle weichselian (the older yoldia sea) and the late weichselian (the younger yoldia sea). modiolaria tumida (hanley 1843) distribution. the british isles and the shetlands, and south to the mediterranean (madsen 1949). in danish waters the species is very common in the limfjord, but also in the other fjords, and it extends into the bælt sea (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. it is a common shallow-water species in danish waters (jensen & spärck 1934), but goes down to 60 m (poppe & goto 1993) associated with tunicates and echinoderms. subfossil finds. the limfjord and vendsyssel regions, holocene. recorded from the north sea region during the eemian. crenella decussata (montagu 1803) distribution.eand w greenland, around iceland, norwaynorthof lofoten and south to the british isles (madsen 1949; petersen 1968). in danish waters the species is found in the kattegat, including the øresund. occurrence. the arctic, subarctic and boreal regions. habitat. from 4 to 200 m deep on all kinds of bottoms (poppe & goto 1993). according to jensen & spärck (1934), most common in danish waters between 15 and 30 m. subfossil finds. recorded in the kattegat during the holocene and from the early and middle weichselian (the older yoldia sea) in the vendsyssel region. adipicola simpsoni (marshall 1900) distribution. from southern iceland and south to portugal and the mediterranean. might be found in danish waters (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. as mentioned by poppe & goto (1993, p. 48): “the species has been collected repeatedly on the skulls of whales, where it lies, attached in the sutures, by its byssus”. subfossil finds. none. order pteroida chlamys islandica (o.f. müller 1776) distribution.seandwgreenland, around iceland, spitsbergen, norway north of lofoten, and south to the shetlands and the orkney islands (rare) (petersen 1968). occurrence. the subarctic and boreal regions. habitat. around iceland one of the most common bivalves present from nearly all localities along the nw, n and e coast, both in the fjords and on the outer part of the shelf, at depth from a few metres to 300 m (madsen 1949). common in danish waters from 10 to 100 m (jensen & spärck 1934) on rocks and gravel bottoms. lives attached to hard substrates with its byssus (poppe & goto 1993). only subfossil finds. recorded from the vendsyssel region during the early and middle weichselian (the olderyoldia sea) and the late weichselian (the younger yoldia sea). aequipecten opercularis (linnaeus 1758) distribution. norway from north of lofoten, and south to the mediterranean (petersen 1968). in danish waters the species extends into the kattegat and øresund, but is not recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone down to 400 m on all types of bottoms except rocky ones (poppe & goto 1993). subfossil finds. the limfjord and vendsyssel regions, holocene. recorded from the north sea region during the eemian. geus bulletin no 3.pmd 28-06-2004, 08:4569 70 chlamys varia (linnaeus 1758) fig. 58 distribution. norway off the lofoten islands, and south to the mediterranean (petersen 1968). in danish waters the species occurs in the limfjord (jensen & spärck 1934) and has been recorded juvenile from the øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. intertidal to 83 m, attached by its byssus (poppe & goto 1993). subfossil finds. the limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen well recorded from the subatlantic. recorded from the north sea during the eemian. delectopecten vitreus (gmelin 1791) fig. 59 distribution. w greenland, w iceland, spitsbergen, norway north of lofoten, and south to the british isles (madsen 1949), and according to poppe & goto (1993) also into the mediterranean. in danish waters from the skagerrak (jensen & spärck 1934). occurrence. the subarctic, boreal and lusitanian regions. habitat. between 30 and 600 m, fixed by its byssus to hard substrates (poppe & goto 1993). the species lives in the deeper part of the skagerrak, from 400 to 600 m, according to jensen & spärck (1934). subfossil finds. the limfjord, holocene. recorded from the eemian in the skagen well. palliolum greenlandicum (sowerby 1842) fig. 60 distribution. e and w greenland, n and e iceland, spitsbergen, and norway north of lofoten (madsen 1949). however, at depths greater than 400 m the species has been found off the faeroes and the british isles. occurrence. the arctic, subarctic and boreal regions. habitat. in the arctic seas living in shallow water from 5 m, but most common between 20 and 70 m on clay bottoms containing stones or shells (ockelmann 1958). only subfossil finds. recorded from the vendsyssel region and the skagen well during the early/middle weichselian (the older yoldia sea). fig. 58. chlamys varia (linnaeus 1758). skagen 3, 32.85–32.90 m b.s., core sample-2. × 20. right valve. mguh 25368. fig. 59. delectopecten cf. vitreus (gmelin 1791). skagen 3, 183.17–183.40 m b.s., lab. no. 9e+05. × 9.6. mguh 25369. fig. 60. palliolum greenlandicum (sowerby 1842). geus collection. east greenland. × 4.8. right valve. mguh 25370. geus bulletin no 3.pmd 28-06-2004, 08:4570 71 palliolum striatum (müller 1776) distribution. from s and w iceland, the faeroes, norway north of lofoten, and south to the mediterranean. the species extends from the north sea and skagerrak into the kattegat, including øresund. occurrence. the boreal and lusitanian regions. habitat. from shallow water around 5 m to more than 800 m deep on all types of bottom (poppe & goto 1993). the vertical range off iceland is indicated to lie between 100 and 260 m (madsen 1949). subfossil finds. the limfjord region, holocene. palliolum tigerinum (müller 1776) distribution. from nw, w and s iceland, the faeroes, norway north of lofoten, and south to morocco. in danish waters from the north sea and skagerrak into the kattegat, including the øresund. occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone down to 400 m, but deeper in the southern part of its range on sandy bottoms (poppe & goto 1993). subfossil finds. the limfjord region, holocene. pecten maximus (linnaeus 1758) distribution. norway off the lofoten islands and south to spain. in danish waters rarely found living in the kattegat and only shells have been recovered from the øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone down to 250 m on sand and gravel bottoms (poppe & goto 1993). subfossil finds. recorded from the vendsyssel region, holocene. pseudamussium septemradiatum (müller 1776) distribution. s iceland, norway from north of the lofoten islands, and south to the mediterranean. in danish waters the species extends from the skagerrak and becomes common in the southern kattegat with finds also in the øresund (jensen & spärck 1934; jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. in general living between 60 and 600 m deep on muddy bottoms (poppe & goto 1993). in danish waters often found between 30 and 60 m. subfossil finds. recorded from the vendsyssel region during the eemian. similipecten similis (laskey 1811) distribution. s and w iceland, the faeroes, norway north of lofoten, and south to the mediterranean (petersen 1968). the species extends into the kattegat from skagerrak. occurrence. the boreal and lusitanian regions. habitat. sublittorally between 4 and 250 m deep on sand and fine gravel bottoms (poppe & goto 1993). in danish waters from 30 to 80 m deep (jensen & spärck 1934). however, the vertical range around iceland is 200–320 m (madsen 1949). subfossil finds. recorded from the vendsyssel area during the eemian. pododesmus patelliformis (linnaeus 1761) distribution. s and w iceland, the faeroes, norway off the lofoten islands, and south to the mediterranean (madsen 1949). the species extends from the north sea into the kattegat, including the øresund, and occurs also in the limfjord (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. intertidal to 50 m deep on gravel or rock bottoms, often attached to shells (poppe & goto 1993). subfossil finds. the limfjord and vendsyssel regions, holocene. pododesmus squama (gmelin 1791) distribution. around the british isles and from danish waters the occurrences in the kattegat and øresund geus bulletin no 3.pmd 28-06-2004, 08:4571 72 are questioned, although larvae occur (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone to 75 m deep on all types of bottoms attached to hard substrates (poppe & goto 1993). subfossil finds. none. anomia ephippium linnaeus 1758 distribution. from the british isles, including the orkney islands, and south to the mediterranean. occurrence. mainly the lusitanian region. habitat. from the intertidal zone down to 150 m on all kinds of hard substrates. only subfossil finds. the limfjord and the vendsyssel region, holocene. heteranomia squamula (linnaeus 1758) fig. 61 distribution. around iceland, the faeroes, norway north of lofoten, and south to the bay of biscay. the species extends into the kattegat, øresund, and occurs in the limfjord (jensen & spärck 1934). occurrence. the subarctic, boreal and lusitanian regions. habitat. from 5 to 110 m deep fixed on hard substrates, but also on algae and crustaceans (poppe & goto 1993). subfossil finds. the kattegat, limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen well recorded from the subatlantic. during the eemian recorded from the bælt sea and the north sea regions. crassostrea gigas (gmelin 1791) this oyster species from the portuguese–spanish region (poppe & goto 1993) has been introduced in 1972 as spat for commercial production (jensen & knudsen 1995) and is not considered here, although mentioned as now part of the danish molluscan fauna. no subfossil records either. ostrea edulis linnaeus 1758 fig. 62a, b distribution. from the southern part of the west coast of norway south to the mediterranean. in danish waters only common in the western part of the limfjord, although stray specimens are found in the northern north sea, skagerrak and northern kattegat (jensen & spärck 1934; jensen & knudsen 1995). fig. 61. heteranomia squamula (linnaeus 1758). skagen 4, 21.0– 21.5 m b.s., lab. no. 346,93. × 20. left valve. mguh 25371. fig. 62. a, b: ostrea edulis linnaeus 1758. skagen 4, 8.0–8.5 m b.s., lab. no. 333,93. × 9.6. left valve of juvenile specimen (exterior and interior, respectively). mguh 25372. geus bulletin no 3.pmd 28-06-2004, 08:4572 73 occurrence. the boreal and lusitanian regions. habitat. from the tidal zone down to 90 m on all types of bottoms. in danish waters the species can be found at depths of 3 to 7 m in the limfjord, but also deeper elsewhere (jensen & spärck 1934). subfossil finds. the bælt sea, kattegat, limfjord, north sea, vendsyssel and skagen regions, holocene. recorded from the subatlantic in the skagen well. finds from the bælt sea, kattegat, and north sea regions during the eemian. limaria hians (gmelin 1791) distribution. norway north of lofoten, the orkney islands and south to the mediterranean (petersen 1968). a few records from the northern and central parts of the kattegat (jensen & knudsen 1995). already jensen & spärck (1934) mentioned that the species then known from the deeper parts of the north sea might occur in danish waters. occurrence. the boreal and lusitanian regions. habitat. lives from the low tide mark zone down to 100 m on coarse sand and gravel bottoms (poppe & goto 1993). subfossil finds. none. limaria loscombi (sowerby 1832) distribution. norway off the lofoten islands, the faeroes and south to the mediterranean (petersen 1968). poppe & goto (1993) mentioned occurrences of l. hians and l. loscombi off iceland which, however, cannot be found in the other literature. in danish waters the species has been found in the kattegat, including the øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. lives from 35 to 100 m deep on fine sand and sand–mud bottoms (poppe & goto 1993). in danish waters it is characteristic in the south-eastern part of kattegat together with pseudamussium septemradiatum (jensen & spärck 1934). subfossil finds. none. limatula subauriculata (montagu 1808) distribution. se and w greenland, n and e iceland, norway north of lofoten, and south to the mediterranean (petersen 1968). in the danish waters only shells have been recorded from the northern kattegat (jensen & knudsen 1995). occurrence. the subarctic, boreal and lusitanian regions. habitat. especially living on the continental shelves. however, records ranges from 4 to 2000 m (poppe & goto 1993). subfossil finds. none. subclass heterodonta order veneroida chama gryphoides linnaeus 1767 this lusitanian species (up to the coasts of portugal) has only one record from danish waters (jensen & knudsen 1995), and this is considered to have been dropped by a ship. therefore it will not be discussed. no subfossil records. lucinella divaricata (linnaeus 1758) distribution. from the english channel and southern part of the north sea south to the mediterranean. occurrence. mainly the lusitanian region. habitat. from the tidal zone down to a depth of 60 m in fine sand and/or mud (poppe & goto 1993). only subfossil finds. the eemian in the bælt sea, kattegat, and north sea regions. lucinoma borealis (linnaeus 1758) distribution. the faeroe islands, norway from off lofoten, and south to the mediterranean (petersen 1968). in danish waters the species occurs in the kattegat, including the øresund, but is not recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. geus bulletin no 3.pmd 28-06-2004, 08:4573 74 habitat. from the intertidal zone down to 500 m deep on gravel bottoms and in pure sand and/or mud (poppe & goto 1993). in danish waters between 20 and 50 m (jensen & spärck 1934). subfossil finds. the limfjord and vendsyssel regions, holocene. myrtea spinifera (montagu 1803) distribution. from norway south to morocco. in danish waters recorded from the northern kattegat (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from 7 to 250 m deep on sand, mud and gravel bottoms (poppe & goto 1993). subfossil finds. none. axinopsida orbiculata (g.o. sars 1878) distribution. e and w greenland, around iceland, the faeroes, and norway north and just south of the lofoten islands (petersen 1968). the species occurs off the north western part of scotland. occurrence. the arctic, subarctic and boreal regions. habitat. from 2 to 50 m deep on sand, clay and mud around iceland (madsen 1949). north of the hebrides occurring at depths down to 900 m (jensen & spärck 1934). only subfossil finds. the vendsyssel region from the early and middle weichselian (the older yoldia sea) and the late weichselian (the younger yoldia sea). thyasira croulinensis (jeffreys 1847) distribution. w greenland, around iceland, the faeroes, norway north of lofoten (petersen 1968), and south to the mediterranean (poppe & goto 1993). in danish waters taken in the north sea and skagerrak. occurrence. the subarctic, boreal and lusitanian regions. habitat. about 40 to 2500 m off the faeroes on gravel and clay (petersen 1968). in danish waters the species is found in the deeper water (jensen & spärck 1934). subfossil finds. none. thyasira equalis (verrill & bush 1898) distribution. it is questioned by nordsieck (1969, p. 79) if t. equalis should be thyasira flexuosa var. rotunda. there are no subfossil finds under the name of t. equalis, so this species will not be considered any further. however, as discussed by ockelmann (1958, p. 100) a species, t. equalis, does occur in the arctic, while t. flexuosa has a boreo-lusitanian main distribution. thyasira flexuosa (montagu 1803) fig. 63 distribution. e and w greenland, spitsbergen, around iceland, the faeroes, norway north of lofoten, and south to the mediterranean. in the danish waters it is very common and extends into the øresund (jensen & knudsen 1995), but is not recorded from the limfjord (petersen 1986a). occurrence. the arctic, subarctic, boreal and lusitanian regions (see comments under t. equalis). habitat. from 10 to 2000 m deep on sand and mud bottoms (poppe & goto 1993). in danish waters from 20 m to around 100 m deep on clay bottoms (jensen & spärck 1934). from the north sea recorded at 30 to 200 m depths on mixed bottom in the trenches around the dogger bank (petersen 1977). subfossil finds. the limfjord, north sea, vendsyssel fig. 63. thyasira flexuosa (montagu 1803). skagen 3, 35.90– 36.00 m b.s., lab. no. 496,93. × 9.6. left valve. mguh 25373. geus bulletin no 3.pmd 28-06-2004, 08:4574 75 and skagen regions, holocene. from the skagen well recorded from the subatlantic. from the north sea during the eemian and in the vendsyssel region recorded from the late weichselian (the younger yoldia sea). thyasira sarsi (philippi 1845) distribution. from novaja semlja along the coast of norway south to the skagerrak region. the species extends into the kattegat, including the øresund (jensen & knudsen 1995). occurrence. the arctic, subarctic and boreal regions. habitat. from 100 m to deep water. subfossil finds. none. leptaxinus ferruginosus (forbes 1844) distribution. w greenland, w iceland, spitsbergen, norway off the lofoten islands, and southwards to madeira (madsen 1949). jensen & spärck (1934) mentioned the species from the deeper part of the skagerrak,but jensen & knudsen (1995) have no further record of this species as belonging to the recent danish fauna. occurrence. the arctic, subarctic, boreal and lusitanian regions. habitat. about 20 m to more than 3000 m. the vertical range off iceland is 320–560 m. so here is another example of tropical submerge (see order mesogastropoda natica affinis). subfossil finds. recorded from the vendsyssel region during the eemian. mysella bidentata (montagu 1803) fig. 64 distribution. around iceland, the faeroes, norway from north of lofoten, and south to west africa (petersen 1968). the species has a common distribution in danish waters, including the limfjord, south to the bælt sea by warnemünde (arntz et al. 1976), and the øresund (jensen & spärck 1934). also found in the north sea (petersen 1977). occurrence. the subarctic, boreal and lusitanian regions. habitat. intertidal zone down to 2500 m, often in commensal association with other animals. in the north sea it is recorded from 20 to 90 m as a commensal on for example, acrocnida brachiata (petersen 1977). subfossil finds. the bælt sea, kattegat, limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen well, records from the preboreal–boreal, subboreal and subatlantic. from the bælt sea, baltic, north sea and vendsyssel regions also recorded from the eemian. mysella tumidula (jeffreys 1867) distribution. this species is recorded by jensen & knudsen (1995) as being part of the danish fauna, although it seems to have a purely lusitanian distribution (poppe & goto 1993). the species has no subfossil occurrence. montacuta substriata (montagu 1803) distribution. w iceland, the faeroes, norway north of lofoten, and south to the mediterranean. in danish waters rarely found in the north sea, skagerrak and kattegat. occurrence. the boreal and lusitanian regions. habitat. off iceland the vertical range is between 31 and 165 m (madsen 1949). in danish waters out to depths of around 700 m (skagerrak) reported as a commensal on spatangus purpureus (jensen & spärck 1934) and from the north sea also on echinocardium flavesens at depths from 30 to 100 m (petersen 1977). subfossil finds. none. fig. 64. mysella bidentata (montagu 1803). skagen 4, 27.0–27.5 m b.s., lab. no. 352,93. × 9.6. to the left a specimen seen from the right, and to the right a left valve. mguh 25374. geus bulletin no 3.pmd 28-06-2004, 08:4575 76 tellimya ferruginosa (montagu 1803) fig. 65 distribution. around iceland, the faeroes, norway north of lofoten, and south to the mediterranean. in danish waters the species extends from the north sea, limfjord, and skagerrak into the kattegat and øresund (jensen & spärck 1934). occurrence. the subarctic, boreal and lusitanian regions. habitat. in general the species is most common just below the tidal zone, which according to poppe & goto (1993) is the preferred habitat of echinocardium with which t. ferruginosa is often associated. however, the species is also found on brissopsis lyrifera or living by itself (jensen & spärck 1934). accordingly, the depth range may vary, around iceland being between 32 and 80 m (madsen 1949). subfossil finds. the limfjord, north sea, vendsyssel and skagen regions, holocene. the records from the skagen well are from the subboreal and subatlantic. furthermore, the species has been recorded from the bælt sea and north sea during the eemian. mysella dawsoni (jeffreys 1864) distribution. petersen (1888, p. 154) mentioned a single find from the limfjord, and the species is mentioned by petersen (1986a) on the basis of the tables on molluscan finds in the limfjord from danmarks fiskeriog havundersøgelser (petersen 1976). however, it is not cited among the recent danish species by jensen & knudsen (1995). the species is mentioned from w greenland (thorson 1951) and south to the mediterranean (poppe & goto 1993), and also from spitsbergen and the west coast of norway (jensen & spärck 1934). occurrence. the arctic, subarctic, boreal and lusitanian regions. habitat. deep-living species, but mentioned from a depth of 5 m by posselt & jensen (1898) and in sandy bottom in west greenland. subfossil finds. none. tellimya tenella (lovén 1846) cited only from the kattegat, including the øresund, by jensen & knudsen (1995). no subfossil finds. will not be considered further. turtonia minuta (fabricius 1780) fig. 66 distribution. w greenland, around iceland, the faeroes, norway north of lofoten, and south to the mediterranean. occurrence. the subarctic, boreal and lusitanian regions. habitat. lives in the tidal zone among plants and algae on rocks (poppe & goto 1993). off iceland in the tidal zone all around the island, but also down to a depth of 50 m (petersen 1968). only subfossil finds. the limfjord, vendsyssel and fig. 65. tellimya ferruginosa (montagu 1803). skagen 4, 27.0– 27.5 m b.s., lab. no. 352,93. × 9.6. to the left interior of a right valve, and to the right a specimen seen from the left. mguh 25375. fig. 66. turtonia minuta (fabricius 1780). skagen 3, 39.85– 40.02 m b.s., lab. no. 711,93. × 40. right valve. mguh 25376. geus bulletin no 3.pmd 28-06-2004, 08:4576 77 skagen regions, holocene. recorded from the subatlantic in the skagen well. lepton nitidum (turton 1822) distribution. from the faeroes, western norway, and south to the mediterranean (jensen & spärck 1934; madsen 1949). only recently recorded from the northern kattegat and øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from 18 to 216 m, commensal on the crustaceans upogebia deltaura and gebia stellata (nordsieck 1969, p. 89). subfossil finds. the limfjord, north sea and vendsyssel regions, holocene. recorded from the eemian in the bælt sea and north sea areas. lepton squamosum (montagu 1803) distribution. from the west coast of norway and south to spain. only shells have been found in danish waters (northern kattegat) mentioned by jensen & knudsen (1995). occurrence. the boreal and lusitanian regions. habitat. from 10 to 120 m deep it lives in and around the burrows of the crustaceans (upogebia deltaura and u. stellata) on mud or gravel bottoms (poppe & goto 1993). subfossil finds. none. devonia perrieri (malard 1904) distribution. from the british isles south to spain. in danish waters recorded from the northern kattegat (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. this species is a commensal on leptosynapta inhaerens (nordsieck 1969, p. 95). subfossil finds. none. kellia suborbicularis (montagu 1803) distribution. s and w iceland, the faeroes, norway off the lofoten islands and south to the mediterranean. in danish waters recorded from the northern kattegat and øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. intertidal to 120 m deep in crevices, shells or in borings made by other species (poppe & goto 1993). subfossil finds. recorded from the vendsyssel region, holocene. potidoma dorkiae (clark 1852) this species has been recorded only in a single find by jensen & knudsen (1995), and there are no subfossil finds from the literature, so it will not be considered further. astarte sulcata (da costa 1778) distribution. se and w greenland, s and w iceland, the faeroes, norway north of lofoten, and south to the mediterranean. the species extends from the north sea into the kattegat, but is not common (jensen & spärck 1934). occurrence. the subarctic, boreal and lusitanian regions. habitat. sublittorally from 20 m and deeper on sand, mud and gravel bottoms. subfossil finds. none. tridonta borealis schumacher 1817 distribution. this species is found in all regions of the north atlantic except the british isles (madsen 1949). however, according to petersen (1968) the species has not been recorded from the faeroes. in the danish waters the species extends from the deeper part of the northern north sea (single finds) into the kattegat, including the øresund and the bælt sea, becoming very common in the baltic (jensen & spärck 1934). occurrence. the arctic, subarctic and boreal regions. geus bulletin no 3.pmd 28-06-2004, 08:4577 78 habitat. from below the tidal zone down to 250 m on mud, sand and gravel bottoms. the species is “a common member of all the zones of the arctic macoma community, and the gomphina fluctuosa community” (ockelmann 1958, p. 83). subfossil finds. the bælt sea and limfjord regions, holocene. recorded from the vendsyssel region both in the early/middle weichselian (not in the skærumhede sequence) and the late weichselian (the older yoldia sea and the younger yoldia sea respectively). tridonta elliptica (brown 1827) distribution. this species is found in all regions of the north atlantic, including the british isles (madsen 1949). petersen (1968) specified the occurrence of this species to be in the clyde sea and off the isle of man only. in the danish waters it occurs in the kattegat, øresund, the bælt sea and the baltic, where it becomes as common as t. borealis (jensen & spärck 1934). occurrence. the arctic, subarctic and boreal regions. considering the occurrences on the west coast of britain it is also in the northern part of the lusitanian region, but here probably mostly in deeper water. habitat. the vertical range for this species is off iceland 6 to 300 m (madsen 1949), and off the east greenland coast it is most often taken between about 5 and 50 m, being abundant locally within the arctic macoma community (ockelmann 1958). subfossil finds. recorded from the north sea during the eemian, and in vendsyssel during the early/middle weichselian, being part of the older yoldia sea found in the skærumhede sequence (jessen et al. 1910). tridonta montagui (dillwyn 1817) distribution. e and w greenland, around iceland, spitsbergen, the faeroes, norway form north of lofoten, and south to the clyde sea and isle of man (petersen 1968), also recorded from the bay of biscay. it is present in the north western part of the north sea (petersen 1977) and common in the kattegat, øresund and extending into the bælt sea (jensen & spärck 1934). occurrence. the arctic, subarctic, boreal and lusitanian regions. habitat. the vertical range off iceland is 7–150 m (madsen 1949) and in the north sea sampled at depths between 40 and 75 m on mixed bottom in the trenches around the dogger bank (petersen 1977, p. 226). subfossil finds. recorded from the north sea region during the eemian and from the vendsyssel area during the early/middle weichselian (the older yoldia sea). acanthocardia echinata (linnaeus 1758) fig. 67 distribution. s and w iceland, norway north of lofoten, and south to the mediterranean (madsen 1949). in danish waters the species extends from the north sea (petersen 1977) into the skagerrak, limfjord, and kattegat regions and the øresund (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. from 4 to 350 m deep on mud, sand and gravel bottoms (poppe & goto 1993). subfossil finds. the kattegat, limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen well recorded from the subboreal and subatlantic. there are eemian records from the bælt sea, kattegat, north sea and vendsyssel regions. parvicardium exiguum (gmelin 1791) distribution. norway north of lofoten and south to the mediterranean (madsen 1949). in danish waters common in bays and fjords, including the limfjord. considering all the variations belonging to the same species it extends into the bælt sea (petersen 1888). occurrence. the boreal and lusitanian regions. fig. 67. acanthocardia echinata (linnaeus 1758). skagen 4, 27.0–27.5 m b.s., lab. no. 352,93. × 20. right valve. mguh 25377. geus bulletin no 3.pmd 28-06-2004, 08:4578 79 habitat. in general occurring from low tide to about 55 m deep (poppe & goto 1993); however, according to rasmussen (1973) the species has its main occurrence along the shores and is associated with vegetation. subfossil finds. the bælt sea, baltic, kattegat, limfjord, north sea and vendsyssel regions, holocene. recorded also from the bælt sea, kattegat, and north sea regions during the eemian. parvicardium hauniense (petersen & russell 1971) distribution. this newly established species has been recorded from recent danish waters, but no subfossil records are at hand. parvicardium ovale (sowerby 1840) distribution. around iceland, norway north of lofoten, and south to the mediterranean (madsen 1949). in danish waters the species is found in all the regions except the baltic extending only to darss (jensen & spärck 1934). occurrence. the subarctic, boreal and lusitanian regions. habitat. according to the icelandic records (madsen 1949) found between 5 to 350 m on bottoms such as mud, sand, clay and shell gravel with stones. subfossil finds. the bælt sea, limfjord, north sea and vendsyssel regions, holocene. from the north sea and vendsyssel regions also recorded from the eemian. parvicardium scabrum (philippi 1844) distribution. from norway north of the lofoten islands, and south to the mediterranean (madsen 1949). in danish waters common in the limfjord (jensen & spärck 1934) and recorded from the kattegat, but questioned, as there may be two separate species (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone to several hundred metres deep on sand, mud and gravel bottoms (poppe & goto 1993). subfossil finds. the bælt sea, kattegat, limfjord, north sea and vendsyssel regions, holocene. recorded from the north sea during the eemian. plagiocardium papillosum (poli 1795) distribution. from the english channel south into the mediterranean (poppe & goto 1993). occurrence. the lusitanian region. habitat. from 1 to 60 m deep on rough sand and gravel bottoms. only subfossil finds. recorded from the eemian in the north sea region. parvicardium minimum (philippi 1836) fig. 68 distribution. s and w iceland, the faeroes, norway north of lofoten, and south to the mediterranean at greater depths. in danish waters common in the deeper part of the skagerrak extending into the kattegat, including the øresund (petersen 1888). the occurrence in the limfjord is questioned (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from 4 to 161 m on mud, sand and gravel bottoms (poppe & goto 1993). however, according to madsen (1949) the species has around iceland only been found at depths of more than 75 m, and the bottoms are recorded as sand with shells and stones or as ooze and clay. in danish waters the species prefers fig. 68. parvicardium minimum (philippi 1836). geus collection. north of skagen, denmark. × 9.6. specimen seen from the right. mguh 25378. geus bulletin no 3.pmd 28-06-2004, 08:4579 80 depths of more than about 30 m, and it occurs at the greatest depth (petersen 1888). subfossil finds. the north sea and skagen regions, holocene. recorded from the skagen well during the preboreal–boreal, the atlantic, subboreal and subatlantic. from the vendsyssel region recorded from the eemian. cerastoderma edule (linnaeus 1758) distribution. norway north of the lofoten islands, and south to the mediterranean (madsen 1949). in danish waters found in all regions (petersen 1888, p. 136 – who already stressed that it is a very variable species). occurrence. the boreal and lusitanian regions. habitat. this is a shallow-water infaunal species – intertidal to few metres deep, but in the baltic occurring also at 20–30 m depths (jensen & spärck 1934) on sandy bottoms. subfossil finds. the bælt sea, baltic, kattegat, limfjord, north sea and vendsyssel regions, holocene. from the eemian recorded in the bælt sea, baltic, kattegat and north sea regions. cerastoderma glaucum (poiret 1789) distribution. from the west coast of norway south to the mediterranean (poppe & goto 1993, pp. 95–96). in danish waters the species extends into the baltic (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. this is a shallow-water species on sand and mud bottoms. however, according to studies on subfossil material (rasmussen 1973, p. 298–302), the associated fauna indicates a tidal estuarine environment for the danish material of an ertebølle age in the isefjord. subfossil finds. the baltic, kattegat and limfjord regions, holocene. however, the species identifications on the subfossil material recorded through time should be taken with some reservation on the basis of the great difficulties connected with recent species identifications. clinocardium ciliatum (fabricius 1780) distribution. w and e greenland, spitsbergen, around iceland and norway north of the lofoten islands. from the faeroes only at depths exceeding 400 m (petersen 1968). occurrence. the arctic, subarctic and boreal (highboreal) regions. habitat. from the tidal zone down to 700 m, off iceland occurring on ooze, mud, clay, sand and mixed bottoms (madsen 1949). mainly found in the arctic macoma community (ockelmann 1958). only subfossil finds. from the vendsyssel region recorded both from the eemian and from the early/middle weichselian (the older yoldia sea). laevicardium crassum (gmelin 1791) distribution. from norway north of the lofoten islands and south to the mediterranean. in danish waters only recorded from the north sea (petersen 1977) and the northern part of the kattegat (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone down to 183 m deep on sand, mud or gravel bottoms. subfossil finds. the vendsyssel region, holocene. from the eemian recorded in the kattegat region. serripes groenlandicus (bruguière 1798) distribution. w and e greenland, spitsbergen, around iceland and norway north of lofoten. from the faeroes only at depths exceeding 400 m (petersen 1968). occurrence. the arctic, subarctic and boreal (highboreal) regions. habitat. from 0 to 1 m to 120 m deep on clay and mud, but also sand and gravel are recorded (madsen 1949). subfossil finds. from the vendsyssel area during the eemian and early/middle weichselian. geus bulletin no 3.pmd 28-06-2004, 08:4580 81 mactra stultorum (linnaeus 1758) fig. 69 distribution. from the british isles and denmark south to the mediterranean (petersen 1968). in the danish waters the species is common in the north sea and skagerrak, extending into the limfjord (petersen 1986a), kattegat and øresund, although here only juveniles are present (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone down to 60 m in clean sand. subfossil finds. the limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen well recorded from the subatlantic. during the eemian recorded from the bælt sea region. lutraria lutraria (linnaeus 1758) distribution. norway off the lofoten islands and south to the mediterranean. since 1990 live specimens have been taken in danish waters near frederikshavn and at the skagerrak coast (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. intertidal down to 100 m, lives at depths up to 35 cm, burrowing in sand, sand–mud or gravel bottoms (poppe & goto 1993). subfossil finds. the limfjord and vendsyssel regions, holocene. spisula elliptica (brown 1827) distribution. around iceland, the faeroes, norway north of lofoten, and south to the english channel and gibraltar. the species occurs in the north sea, extending into the limfjord and kattegat and the øresund (jensen & knudsen 1995). occurrence. the subarctic, boreal and lusitanian regions. habitat. between 20 and 200 m deep in mud, sand and gravel bottoms. subfossil finds. the limfjord, north sea and vendsyssel regions, holocene. from the bælt sea region recorded during the eemian. spisula solida (linnaeus 1758) distribution. s and w iceland, and south to the mediterranean (petersen 1968). in danish waters very common offshore from the west coast of jylland (jensen & knudsen 1995 – as recorded by petersen 1977, fig. 25). the species extends into the limfjord (petersen 1986a) and kattegat, including the øresund and the bælt sea regions (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone down to 100 m on sandy bottoms. in the north sea found at 15–40 m depths in coarse sand (petersen 1977). subfossil finds. the limfjord, north sea and vendsyssel regions, holocene. recorded from the kattegat during the eemian. spisula subtruncata (da costa 1778) fig. 70a, b distribution. norway north of the lofoten islands, and south to the mediterranean (petersen 1968). in danish waters the species extends from the north sea into the kattegat, where it is also common. the species has been recorded both from the øresund and the bælt sea (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone down to 200 m deep fig. 69. mactra stultorum (linnaeus 1758). skagen 4, 15.0–15.5 m b.s., lab. no. 340,93. × 4.8. hinges of two right valves. mguh 25379. geus bulletin no 3.pmd 28-06-2004, 08:4581 82 in mud and sand (poppe & goto 1993). in danish waters common at depths out to 20–30 m in sand (jensen & spärck 1934). subfossil finds. the bælt sea, limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen well recorded from the atlantic, subboreal and subatlantic; in the subatlantic occurring in huge quantities. from the eemian recorded from the bælt sea, baltic and kattegat regions. solecurtus chamasolen (da costa 1778) distribution. norway off the lofoten islands, and south to the mediterranean. in the danish waters shells have been found near frederikshavn (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from 5 to 400 m deep on muddy bottoms. subfossil finds. none. solecurtus scopula (turton 1822) distribution. from the british isles and south to the mediterranean. shells recorded from the northern kattegat (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. mainly offshore to 110 m deep in muddy sand and on clean gravel bottoms (poppe & goto 1993). subfossil finds. none. ensis americanus gould 1870 distribution. at present occurring down to the øresund. however, the species has recently accidentally been transported to western europe from the east coast of north america (jensen & knudsen 1995), so it will not be further considered. ensis arcuatus (jeffreys 1865) distribution. the faeroes, the british isles and south to spain. in danish waters shells have been recorded from the kattegat region, including the øresund. occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone to 40 m deep in sand and gravel bottoms. subfossil finds. none. ensis ensis (linnaeus 1758) distribution. the faeroes, norway north of lofoten, and south to the mediterranean (petersen 1968). the species extends from the north sea east to the limfjord and kattegat regions, including the øresund (jensen & spärck 1934; jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone to about 80 m deep, burrowing in fine sand (poppe & goto 1993). from the north sea mainly recorded from depths of 20–30 m on the dogger bank and along the west coast of jylland (petersen 1977, fig. 40). fig. 70. a: spisula subtruncata (da costa 1778). skagen 4, 15.0–15.5 m b.s., lab. no. 340,93. × 4.8. right valve interior. mguh 25380. b: spisula subtruncata (da costa 1778). skagen 4, 15.0–15.5 m b.s., lab. no. 340,93. × 20. serrated surfaces of paired right valve laterals. mguh 25380. geus bulletin no 3.pmd 28-06-2004, 08:4582 83 subfossil finds. the limfjord, north sea and vendsyssel areas, holocene. during the eemian recorded from the bælt sea, north sea and the vendsyssel regions. ensis siliqua (linnaeus 1758) distribution. norway from north of lofoten, and south to the mediterranean. in danish waters, the species is recorded from the north sea, kattegat and øresund (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone to about 70 m deep in fine sand. subfossil finds. none. phaxas pellucidus (pennant 1777) fig. 71 distribution. norway off the lofoten islands, and south to the mediterranean (madsen 1949). in danish waters the species extends into the limfjord and kattegat, øresund and bælt sea regions (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. found offshore between 4 and 150 m deep in sand, mud and gravel bottoms (poppe & goto 1993). in the north sea it is abundant in the whole area, mainly from depths of 30–50 m (petersen 1977). subfossil finds. the limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen well recorded from the subboreal and subatlantic. during the eemian found in the north sea and vendsyssel regions. angulus tenuis (da costa 1778) fig. 72 distribution. norway off the lofoten islands, and south to the mediterranean (madsen 1949). in danish waters the species extends from the north sea, where it is common in shallow waters, into the kattegat and limfjord regions (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. common in shallow water down to 10–20 m deep in fine sand. subfossil finds. the bælt sea, limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen well recorded from the subatlantic. during the eemian found in the bælt sea and north sea regions. arcopagia crassa (pennant 1778) distribution. from norway south to west africa. only one record from danish waters other than empty shells (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone down to 150 m deep in sand, mud and shell gravel bottoms (poppe & goto 1993). subfossil finds. none. fig. 71. phaxas pellucidus (pennant 1777). skagen 4, 27.0–27.5 m b.s., lab. no. 352,93. × 9.6. view of the inside of a left valve above and a right valve below (valves not paired). mguh 25381. fig. 72. angulus tenuis (da costa 1778). skagen 4, 6.0–6.5 m b.s., lab. no. 331,93. × 4.8. view of the inside of a right valve. mguh 25382. geus bulletin no 3.pmd 28-06-2004, 08:4583 84 tellina donacina linnaeus 1758 distribution. from the shetlands over the british isles and south into the mediterranean (poppe & goto 1993). the species occurs in the southern north sea and has been recorded off edinburgh (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. from the low-tide mark to 200 m deep in sand, mud and gravel bottoms. only subfossil finds. recorded from the north sea during the eemian. fabulina fabula (gmelin 1791) fig. 73a, b distribution. norway of the lofoten islands, and south to the mediterranean (madsen 1949). the species extends into danish waters from the north sea to the kattegat and limfjord regions (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone to 50 m deep in fine sand. in danish waters common on sand between 5– 25 m, but it has been taken at depths of 40 m (jensen & spärck 1934). subfossil finds. the limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen well recorded from the subatlantic. from the eemian recorded from the north sea. tellina pygmaea (lovén 1846) fig. 74 distribution. the faeroes, norway north of lofoten, and south to the mediterranean (madsen 1949). the species extends from the north sea into the kattegat and øresund (jensen & knudsen 1995). uncertain in the records from the limfjord (collin 1884, p. 113). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone to depths of 100 m. in the north sea found at depths of 30–50 m on hard bottoms (petersen 1977). subfossil finds. from the skagen well recorded from the subatlantic. gastrana fragilis (linnaeus 1758) distribution. from the british isles including the shetlands, and south to the mediterranean. fig. 73. a: fabulina fabula (gmelin 1791). skagen 4, 6.0–6.5 m b.s., lab. no. 331,93. × 4.8. view of the inside of a right valve. mguh 25383. b: × 4.8. right valve exterior with diagonal lines running from the upper right to the lower left superimposed upon sculpture of concentric lines. mguh 25383. fig. 74. tellina pygmaea (lovén 1846). skagen 4, 27.0–27.5 m b.s., lab. no. 352,93. × 40. right valve. mguh 25384. geus bulletin no 3.pmd 28-06-2004, 08:4584 85 occurrence. mainly the lusitanian region. habitat. from below the tidal zone to a depth of 750 m in sand, mud and gravel bottoms (poppe & goto 1993). subfossil finds. recorded from the bælt sea, kattegat and the north sea during the eemian. macoma balthica (linnaeus 1758) distribution. w greenland, norway from north of lofoten, and south to the british isles (petersen 1968) and spain (poppe & goto 1993). the species is found in all the regions and extends far into the baltic, but it is not common in the north sea region from blåvandshuk and north to skagen (jensen & spärck 1934). occurrence. the subarctic, boreal and lusitanian regions. habitat. this is a shallow-water species, but in the baltic occurs also at depths of more than 50 m on soft bottoms (muus 1967, p. 163). the species is also the characterising animal of the petersen macoma balthica community so often found in the inner danish waters. however, muus (op. cit., pp. 215–217) discussed the problem concerning this community in further detail and concluded that petersen’s community can be considered a serviceable way of giving a brief description of a faunal region for other marine biologists. subfossil finds. the bælt sea, baltic, kattegat, limfjord, north sea and vendsyssel regions, holocene. recorded from the bælt sea, baltic, and the north sea during the eemian, and from the vendsyssel area during the late weichselian. macoma calcarea (gmelin 1791) distribution. w and e greenland, spitsbergen, around iceland, the faeroes, norway from north of lofoten (madsen 1949), and south into the north sea, kattegat, bælt sea and baltic (jensen & spärck 1934). occurrence. the arctic, subarctic and boreal regions. habitat. from the intertidal zone down to several hundred metres in the southern part of the distribution area. the species is the characteristic animal of the arctic macoma community (thorson 1957). subfossil finds. from the vendsyssel region, holocene. recorded from the baltic and the vendsyssel regions during the eemian, from the kattegat and vendsyssel regions during the early/middle weichselian (the older yoldia sea stage), and finally from the vendsyssel region also during the late weichselian (the younger yoldia sea stage). macoma torelli (steenstrup) jensen 1904 distribution. e and w greenland and spitsbergen. according to ockelmann (1958) this species is regarded as having a high-arctic main distribution and being restricted to the n atlantic sector. occurrence. the arctic region. habitat. recorded rarely from greenland sublittorally out to 90 m deep on clay and gravel (ockelmann 1958). subfossil finds. recorded from the vendsyssel region during the late weichselian (the younger yoldia sea). macoma loveni (steenstrup) jensen 1904 distribution. w and e greenland and spitsbergen. according to ockelmann (1958), the main distribution is high-arctic with low-arctic outposts. occurrence. the arctic and subarctic regions. habitat. at east greenland the vertical range of the species has been recorded from 4–5 m to 207 m, and the species has been taken from various types of bottoms: clay, mud, sand, gravel and stony ones (ockelmann 1958). subfossil finds. the vendsyssel region during the late weichselian (the younger yoldia sea). donax vittatus (da costa 1778) fig. 75 distribution. norway off the lofoten islands, and south to the mediterranean (madsen 1949). this species is found on the southern part of the dogger bank and along the west coast of jylland (petersen 1977) but not in the inner danish waters (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. geus bulletin no 3.pmd 28-06-2004, 08:4585 86 habitat. from the tidal zone down to 20 m in clean sand. one of the few molluscan species well suited to live in the coastal zone of sandy beaches. subfossil finds. the limfjord on old beach ridges facing the skagerrak (petersen 1976), north sea, vendsyssel and skagen regions, holocene. from the skagen well recorded from the subatlantic. found in the north sea during the eemian. gari depressa (pennant 1777) distribution. norway off the lofoten islands, and south to the mediterranean. occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone down to about 50 m in sand, mud and gravel bottoms. subfossil finds. the vendsyssel region, holocene. gari fervensis (gmelin 1791) fig. 76 distribution. w greenland, the faeroes, norway north of lofoten, and south to the mediterranean (petersen 1968). in the north sea the species is most common on the dogger bank (petersen 1977) and extends through the skagerrak into the kattegat and øresund (jensen & spärck 1934). occurrence. the subarctic, boreal and lusitanian regions. habitat. from the tidal zone to a depth of 110 m in coarse sand and shell gravel bottoms (poppe & goto 1993). in danish waters on mixed bottoms and sand at depths of 15–40 m (jensen & spärck 1934). subfossil finds. the limfjord, vendsyssel and skagen regions, holocene. recorded from the subatlantic in the skagen well. gari tellinella (lamarck 1818) distribution. s and w iceland, the faeroes, norway north of lofoten, and south to the mediterranean. the species has been taken from the northern kattegat (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone to a depth of 460 m in coarse sand, gravel and stone bottoms. subfossil finds. none. scrobicularia plana (da costa 1778) distribution. norway off the lofoten islands, and south to the mediterranean (madsen 1949). the species extends into the bælt sea at kiel and warnemünde (jensen & spärck 1934), and is also recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone to about 30 m in clay or muddy bottoms, often in estuaries (poppe & goto 1993). jensen & knudsen (1995) point to the occurrences in the wadden sea and to the sensitivity to severe winters, living in such shallow-water environments. fig. 75. donax vittatus (da costa 1778). geus collection. holland. × 4.8. view of the inside of a right valve. mguh 25385. fig. 76. gari fervensis (gmelin 1791). skagen 4, 23.0–23.5 m b.s., lab. no. 348,93. × 4.8. right valve. mguh 25386. geus bulletin no 3.pmd 28-06-2004, 08:4586 87 subfossil finds. the bælt sea, baltic, kattegat, limfjord, north sea and vendsyssel regions, holocene. during the eemian recorded from the bælt sea, baltic, kattegat, and north sea regions. abra alba (wood 1802) fig. 77 distribution. norway off the lofoten islands, and south to the mediterranean (madsen 1949). in danish waters it is found most abundantly in the shallow-water parts of the north sea (petersen 1977) and extends through the skagerrak, limfjord and kattegat and øresund into inner danish waters such as the bælt sea and the baltic to neustadt, where it is the typical bottom animal (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. from the infralittoral zone to a depth of 65 m in sand, mud or muddy gravel (poppe & goto 1993). in danish waters common at depths of 3–8 m and out to 15–20 m in soft bottoms (jensen & spärck 1934). subfossil finds. the bælt sea, kattegat, limfjord, north sea, vendsyssel and skagen regions, holocene. in the skagen well recorded from the subboreal and subatlantic. during the eemian recorded from the bælt sea, north sea and vendsyssel regions. abra nitida (müller 1776) fig. 78 distribution. recorded from s and w iceland, the faeroes, norway north of lofoten, and south to the mediterranean (madsen 1949). in danish waters recorded from the north sea, skagerrak, limfjord and kattegat (jensen & spärck 1934), including the øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. mainly offshore to depths of 200 m in sandy mud, mud or gravel bottoms (poppe & goto 1993). however, in danish waters such as the north sea on soft to mixed bottoms at depths of 40–70 m (petersen 1977) and in the skagerrak at depths of 100–300 m extending into the kattegat, including the øresund, on soft bottoms (jensen & spärck 1934). subfossil finds. the limfjord, north sea, vendsyssel and skagen regions, holocene. in the skagen well recorded from the subboreal and subatlantic. during the eemian recorded from the vendsyssel region. abra prismatica (montagu 1803) fig. 79 distribution. s and w iceland, the faeroes, norway north of lofoten, and south to the mediterranean (madsen 1949). in the danish waters the species extends from the north sea and skagerrak into the kattegat and øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from the infralittoral zone to 400 m deep in sand or muddy sand (poppe & goto 1993). in the north sea most abundant at depths deeper than 50 m on mixed bottoms (petersen 1977). subfossil finds. the limfjord, vendsyssel and skagen regions, holocene. from the skagen well recorded from the subboreal and subatlantic. during the eemian refig. 77. abra alba (wood 1802). geus collection. storebælt, denmark. × 4.8. specimen seen from the left. mguh 25387. fig. 78. abra nitida (müller 1776). geus collection. hellebæk, denmark. × 4.8. specimen seen from the left. mguh 25388. geus bulletin no 3.pmd 28-06-2004, 08:4587 88 corded from the bælt sea, kattegat, north sea, and vendsyssel regions. abra segmentum (récluz 1843) distribution. from the west coast of france into the mediterranean. occurrence. the lusitanian region. habitat. in the infralittoral zone in sandy mud (poppe & goto 1993). it seems to be connected with the shallow-water environment, also with brackish water (jensen & spärck 1934). only subfossil finds. recorded from the bælt sea and the north sea during the eemian. arctica islandica (linnaeus 1767) fig. 80 distribution. around iceland, the faeroes, norway north of lofoten, and south to the bay of biscay. in the danish waters, including the limfjord, the species extends from the north sea and skagerrak as far as the baltic (to bornholm) (jensen & spärck 1934). occurrence. the subarctic, boreal and (lusitanian) regions. however, the species tends to live more deeply in the southern part of its range (poppe & goto 1993). habitat. intertidal to 482 m in mud, sand or gravel bottoms. in the north sea mainly from depths deeper than 40 m and from mixed bottoms (petersen 1977). in inner danish waters often at depths from 10–15 to 50–60 on clay or clayey bottoms (jensen & spärck 1934). according to badarsson (1920), fishermen say that the species occurs in very shallow water, just below the low-water mark, in winter living deeply burrowed in the substrate, but in the summer often lying in abundance on the bottom. subfossil finds. the kattegat, limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen well recorded in the subatlantic. in the bælt sea, baltic and north sea (there are (single find) records from the eemian and the vendsyssel region during the late weichselian (the younger yoldia sea). kelliella miliaris (philippi 1844) fig. 81 distribution. from norway off the lofoten islands, and south to the mediterranean. from danish waters recorded from the skagerrak. according to jensen & spärck (1934), very common in the deeper part of the skagerrak. jensen & knudsen (1995) mentioned a single finding from the southern kattegat. fig. 79. abra prismatica (montagu 1803). geus collection. iceland. × 4.8. specimen seen from the left. mguh 25389. fig. 80. arctica islandica (linnaeus 1767). geus collection. læsø rende, denmark. × 4.8. juvenile specimen, beaks directed forwards. mguh 25390. fig. 81. kelliella miliaris (philippi 1844). skagen 3, 185.04–185.06 m b.s., core sample-102. × 20. two specimens seen from the left. inside the corroded specimen to the left, pyrite is seen. mguh 25391. geus bulletin no 3.pmd 28-06-2004, 08:4588 89 occurrence. the boreal and lusitanian regions. habitat. from 134 to 700 m deep (nordsieck 1969). subfossil finds. recorded from the eemian in the vendsyssel and skagen regions. glossus humanus (linnaeus 1758) distribution. s and w iceland, norway off the lofoten islands, and south to the mediterranean. in danish waters recorded from the north sea and kattegat, where shells are common (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. offshore beyond 7 m on bottoms of sand, sandy mud or soft mud (tebble 1966). subfossil finds. none. chamelea striatula (da costa 1778) fig. 82 distribution. the faeroes, norway north of lofoten, and south to the mediterranean (madsen 1949). in danish waters the species extends from the north sea, limfjord and skagerrak into the kattegat and øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from the infralittoral zone to 55 m deep on sand and mud bottoms (poppe & goto 1993). in the north sea mainly from 20 to 40 m on sand to mixed bottoms (petersen 1977). in the skagerrak at depths of more than 100 m, but in the kattegat at depths of less than 50 m, since the sand bottom is not to be found at deeper levels (jensen & spärck 1934). subfossil finds. the limfjord, north sea, vendsyssel and skagen areas, holocene. from the skagen well recorded from the subboreal and subatlantic. during the eemian recorded from the bælt sea, north sea and vendsyssel regions. clausinella fasciata (da costa 1778) distribution. the faeroes, norway off the lofoten islands, and south to the mediterranean. from the north sea the species extends into the kattegat (petersen 1888). occurrence. the boreal and lusitanian regions. habitat. from 4 to 110 m deep in sand, mud and gravel bottoms (poppe & goto 1993). in the north sea in hard sand (petersen 1977), and in the kattegat in gravel and sand between 15 and 30 m (petersen 1888, p. 143). subfossil finds. the limfjord, north sea and vendsyssel areas, holocene. paphia aurea (gmelin 1791) distribution. norway off the lofoten islands, and south to the mediterranean. the finds closest to danish waters are from southern norway (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone down to 36 m deep in sand, mud and gravel bottoms. only subfossil finds. the bælt sea, kattegat, limfjord, north sea and vendsyssel regions, holocene. recorded from the kattegat region during the eemian. paphia aurea senescens (cocconi 1873) distribution. the only one of our quaternary molluscs which does not live at present. also found in quaternary deposits in italy (jensen & spärck 1934). regarded as a lusitanian species according to nordmann (1913). however, poppe & goto (1993) regard the fossil valves found, for example along the coast of the netherlands and belgium, as a subspecies in which the differences fig. 82. chamelea striatula (da costa 1778). skagen 4, 21.0– 21.5 m b.s., lab. no. 346,93. × 4.8. left valve. mguh 25392. geus bulletin no 3.pmd 28-06-2004, 08:4589 90 from extant ones are minimal, and they propose that the relationship between the fossil and recent shells be restudied. from nordmann (1913) and cerulli-irelli (1908) it seems right that tapes senescens doederl. and tapes aureus var. eemiensis are identical. but as the relationship between t. senescens and t. aureus sensu stricto at the time of nordmann (1913) was not clear, the position as a not extant subspecies given by poppe and goto is followed here. only subfossil finds. the bælt sea, kattegat, and north sea regions during the eemian. tapes decussatus (linnaeus 1758) distribution. norway off the lofoten islands, and south to the mediterranean. closest to danish waters the species occurs off western norway (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. from the tidal zone to a depth of few metres in sand or muddy-gravel bottoms. only subfossil finds. the kattegat, limfjord, north sea and vendsyssel regions, holocene. recorded from the north sea region during the eemian. timoclea ovata (pennant 1777) fig. 83a, b distribution. s and w iceland, the faeroes, norway north of lofoten, and south to the mediterranean (madsen 1949). in danish waters the species extends from the north sea into the kattegat and øresund (jensen & knudsen 1995), but has not been recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. at depths between 4 and 200 m on all types of bottoms (poppe & goto 1993). in the north sea usually deeper than 50 m and on soft bottoms (petersen 1977). subfossil finds. the limfjord, north sea, vendsyssel and skagen regions, holocene. in the skagen well recorded from the subatlantic. during the eemian found in the bælt sea, kattegat and north sea regions. venerupis rhomboides (pennant 1777) distribution. the faeroes, norway off the lofoten islands (madsen (1949) does not mention any norwegian occurrence), and south to the mediterranean (petersen 1968). one pair of united valves recorded from the northern kattegat (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone to 180 m deep (poppe & goto 1993) in gravel and mud bottoms. subfossil finds. the limfjord and vendsyssel regions, holocene. venerupis pullastra (montagu 1803) distribution. norway from north of the lofoten islands, and south to the mediterranean (madsen 1949). from danish waters recorded from the limfjord and kattefig. 83. a: timoclea ovata (pennant 1777). skagen 4, 27.0–27.5 m b.s., lab. no. 352,93. × 40. left valve. mguh 25393. b: timoclea ovata (pennant 1777). geus collection. herthas flak, denmark. × 4.8. right valve. mguh 25394. geus bulletin no 3.pmd 28-06-2004, 08:4590 91 gat regions (jensen & spärck 1934) and the øresund (jensen & knudsen 1995). the species has a common occurrence in the isefjord (rasmussen 1973). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone to 40 m deep in hard sand and muddy gravel (poppe & goto 1993). rasmussen (1973, p. 303): considers “its present common occurrence there [in the isefjord] to be a result of the disappearance of the zostera since 1933–45, as the lack of a continuous vegetation caused a change in the bottom conditions to the benefit of venerupis pullastra. undoubtedly the species has lived in interior danish waters since the stone age, being, however, rare in recent times up to 1933–34”. subfossil finds. the bælt sea, kattegat, limfjord, north sea and vendsyssel regions, holocene. recorded from the north sea region during the eemian. dosinia exoleta (linnaeus 1758) distribution. norway off lofoten, and south to the mediterranean. the species extends from the north sea (petersen 1977) into the northern kattegat (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone to 73 m deep, burrowing deeply in sand, mud and gravel bottoms (poppe & goto 1993). in the north sea found at depths of 30– 100 m, on hard bottoms around the dogger bank (petersen 1977). subfossil finds. the north sea and vendsyssel regions, holocene. dosinia lincta (montagu 1803) distribution. s and w iceland, the faeroes, norway north of lofoten, and south to the mediterranean (petersen 1968). the species extends from the north sea into the kattegat and øresund (petersen 1888). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone to 200 m deep in pure sand and fine gravel bottoms (poppe & goto 1993). in the north sea sampled at 10–100 m on soft to mixed bottoms (petersen 1977). from the kattegat recorded on mixed bottoms between 18 and 56 m (petersen 1888). subfossil finds. the north sea and vendsyssel regions, holocene. recorded from the bælt sea and the north sea during the eemian. gouldia minima (montagu 1803) distribution. the faeroes, norway off the lofoten islands, and south to the mediterranean (petersen 1968). occurrence. the boreal and lusitanian regions. habitat. from below the tidal zone to depths of over 200 m in sand, mud and fine gravel bottoms (poppe & goto 1993). only subfossil finds. recorded from the bælt sea and the north sea during the eemian. petricola pholadiformis (lamarck 1822) distribution. this species in an immigrant introduced in to europe, probably with oysters, at the end of the last century (poppe & goto 1993). according to jensen & knudsen (1995), the species occurred in 1906 in the wadden sea; the skagerrak 1905; the kattegat 1931; and the bælt sea 1943. it will not be considered any further in this work. mysia undata (pennant 1777) distribution. the faeroes, norway north of the lofoten islands, and south to the mediterranean (petersen 1968). in danish waters the species extends from the north sea (petersen 1977) into the kattegat and øresund (jensen & knudsen 1995). it is not recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from below the tidal zone to depths of 55 m in muddy sand and gravel bottoms (poppe & goto 1993). in the north sea found at depths of 40–70 m on soft to mixed bottoms (petersen 1977). subfossil finds. the limfjord and vendsyssel regions, holocene. during the eemian recorded from the north sea region. geus bulletin no 3.pmd 28-06-2004, 08:4591 92 order myoida mya arenaria linnaeus 1758 distribution. norway from north of the lofoten islands, south to the british isles (petersen 1968). the species extends from the north sea into the limfjord and the inner danish waters, including the baltic (jensen & spärck 1934). it is a late immigrant, known from europe in the plio-pleistocene (strauch 1972, pp. 135– 137) having been transferred from north america by man, presumably the vikings, and dated back to the 13th century, i.e. well before columbus (petersen et al. 1992b). occurrence. mainly in the boreal region, but with new finds further to the south on the east coast of north america (rasmussen & heard 1994). habitat. from the tidal zone down to 6–7 m deep in sandy bottoms, the species avoids high-energy coastal environments (jensen & spärck 1934). subfossil finds. the records from the bælt sea and vendsyssel might be of recent dates; only the occurrences at jerup halfway up to the skagen spit have been dated and included in the skagen area from the subatlantic. mya truncata linnaeus 1758 distribution. w and e greenland, spitsbergen, around iceland, the faeroes, norway from north of lofoten, and south to the bay of biscay (madsen 1949; poppe & goto 1993). the species extends from the north sea into the limfjord and the inner danish waters as far as the bælt sea (kiel and warnemünde) (jensen & spärck 1934). occurrence. the arctic, subarctic, boreal and lusitanian regions. habitat. from the intertidal zone down to about 75 m deep (poppe & goto 1993). however, in danish waters often at depths between 10 and 20 m (jensen & spärck 1934). in the north sea found at depths of 37– 70 m on soft mixed bottoms (petersen 1977). in east greenland the species belongs to the arctic macoma community (ockelmann 1958). however, according to jensen (1900) the typical mya truncata is not found in thearctic (g.h.petersen,personalcommunication1998). subfossil finds. the bælt sea, kattegat, limfjord, north sea and vendsyssel regions, holocene. recorded from the eemian in the bælt sea, baltic, north sea and vendsyssel areas. from the early/middle weichselian in the kattegat and vendsyssel regions, and from the vendsyssel region also in the late weichselian. corbula gibba (olivi 1792) fig. 84 distribution. norway north of lofoten, and south to the mediterranean (petersen 1968). the species extends from the north sea into the inner danish waters, including the limfjord, as far as the westernmost part of the baltic region at møn (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. from the low intertidal zone to 250 m deep anchored by a byssus on silty sand and muddy-gravel bottoms (poppe & goto 1993). however, in danish waters rarely at depths of more than about 50 m (jensen & spärck 1934). in the north sea sampled at 35–50 m depths on mixed bottoms (petersen 1977). subfossil finds. the bælt sea, baltic, kattegat, limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen well recorded from the subboreal and subatlantic. during the eemian recorded from the bælt sea, baltic, kattegat, north sea and vendsyssel regions. fig. 84. corbula gibba (olivi 1792). skagen 3, 37.00–37.25 m b.s., lab. no. 710,93. × 9.6. to the upper left a specimen seen from the left, and to the right a view of the inside of a right valve. mguh 25395. geus bulletin no 3.pmd 28-06-2004, 08:4592 93 hiatella arctica (linnaeus 1758) fig. 85 distribution. w and e greenland, spitsbergen, around iceland, the faeroes, norway from north of lofoten, and south to the mediterranean (madsen 1949). following jensen & spärck (1934) who regard the records of hiatella as one species, it is widely spread but not always common in all the danish waters, it extends to kiel in the bælt sea. occurrence. the arctic, subarctic, boreal and lusitanian regions. habitat. “from the intertidal zone down to almost 1400 m fixed by its byssus on or in all kinds of substrate on all types of bottoms. also found in holes previously bored by other species” (poppe & goto 1993, p. 130). however, petersen (1977, p. 228) states: “both the systematics and nomenclature are insufficiently investigated for this genus”. so with regard to the subfossil finds, the questions on species are even more difficult, as seen from petersen (1986b, figs 2, 3), where forms with different habitat such as hiatella cf. byssifera are found fixed on a stone taken as a grab sample in kejser franz josephs fjord, east greenland, and as traces of hiatella arctica in the saxicava sand of late weichselian age in vendsyssel. however, here símonarson et al. (1998) is followed, relating the more widely used and less specific name hiatella arctica. subfossil finds. the bælt sea, limfjord, north sea, vendsyssel and skagen areas, holocene. from the skagen well records from the subboreal and subatlantic. from the eemian recorded from the kattegat, north sea, vendsyssel and skagen regions. in the kattegat and vendsyssel regions finds from the early/middle weichselian (the older yoldia sea) and in the vendsyssel region from the late weichselian (the younger yoldia sea). hiatella rugosa (linnaeus 1758) distribution. jensen & knudsen (1995) include this as a separate species and take it as part of the recent danish shell-bearing fauna. records from the literature on subfossil finds are therefore considered here. according to poppe & goto (1993), found from norway south to the mediterranean. ockelmann (1958, p. 135 ff.), discussing the hiatella taxonomy at some length, concludes that reservations must be made as to future separations of the hiatella forms (h. arctica incl. of h. gallicana and h. pholadis) occurring in the northern hemisphere into valid species. occurrence. the arctic, subarctic, boreal and lusitanian regions. subfossil finds. the kattegat and vendsyssel regions, holocene. saxicavella jeffreysi winckworth 1930 fig. 86 distribution. norway off the lofoten islands, and south to the mediterranean (madsen 1949). the species extends into the kattegat and øresund, although rare (jensen & knudsen 1995). it has been recorded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. offshore between 7 and 240 m deep in sand, mud or gravel bottoms (poppe & goto 1993). in the fig. 85. hiatella arctica (linnaeus 1758). skagen 3, 37.00–37.25 m b.s., lab. no. 710,93. × 20. left valve. mguh 25396. fig. 86. saxicavella jeffreysi winckworth 1930. skagen 4, 20.0– 20.5 m b.s., lab. no. 345,93. × 4.8. right valve. mguh 25397. geus bulletin no 3.pmd 28-06-2004, 08:4593 94 danish waters often at depths between 25–50 m (jensen & spärck 1934). subfossil finds. the limfjord, north sea, vendsyssel and skagen regions, holocene. in the skagen well recorded from the subatlantic. during the eemian recorded from the north sea region. panomya arctica (lamarck 1818) distribution. s and w iceland, norway from north of lofoten, and south to the british isles and denmark (madsen 1949). however, petersen (1968) refers to empty shells from n iceland and occurrences in the clyde sea and off the isle of man. in danish waters only once taken alive near skagen, otherwise shells only, but found as far south as øresund (jensen & knudsen 1995). occurrence. mainly boreal, but outposts into the subarctic (subfossil?) and lusitanian regions (strauch 1972). habitat. from the intertidal zone down to 300 m buried in mud or sand (poppe & goto 1993). subfossil finds. none. barnea candida (linnaeus 1758) fig. 87 distribution. norway off the lofoten islands, and south to the mediterranean (madsen 1949). in the danish waters, including the limfjord, the species extends into the bælt sea as far as kiel (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. from the low intertidal zone to about 30 m deep, the species bores in semi-hard substrates such as clay (poppe & goto 1993). subfossil finds. the bælt sea, limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen well recorded from the subatlantic. during the eemian found in the bælt sea and north sea regions. pholas dactylus linnaeus 1758 fig. 88a, b distribution. norway off the lofoten islands, and south to the mediterranean (madsen 1949). in danish waters only found to frederikshavn and the limfjord (petersen 1986a; jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone to a depth of 10 m boring in different substrates, preferring clay bottoms (poppe & goto 1993). subfossil finds. the limfjord, vendsyssel and skagen regions, holocene. from the skagen well recorded from the subatlantic. zirfaea crispata (linnaeus 1758) distribution. around iceland, norway from north of the lofoten islands, and south to the bay of biscay (madsen 1949). in the danish waters, including the limfjord, extending into the bælt sea as far as kiel (jensen & spärck 1934). occurrence. mainly boreal with outposts into the lusitanian region. habitat. from the low tide line to about 7 m deep, boring in semi-hard substrates (poppe & goto 1993). in the danish waters the species has a wide extension, depending on the bottom substrates (jensen & spärck 1934): in the north sea and skagerrak peat and chalk; in the limfjord cementstone, mo-clay, chalk, peat and clay; and the bælt sea clayey bottoms. subfossil finds. the bælt sea, limfjord, north sea and vendsyssel regions, holocene. from the bælt sea region recorded from the eemian and from the vendsyssel region during the late weichselian (the younger yoldia sea). fig. 87. barnea candida (linnaeus 1758). geus collection. off rubjerg knude, denmark. × 4.8. left valve. mguh 25398. geus bulletin no 3.pmd 28-06-2004, 08:4594 95 xylophaga dorsalis turton 1822 distribution. s and w iceland, norway from north of lofoten, and south to the mediterranean. recorded from the øresund region (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. according to madsen (1949) the occurrences at iceland have been from depths between 140–230 m found in sunken pieces of wood. subfossil finds. none. teredo navalis linnaeus 1758 distribution. norway off the lofoten islands (petersen 1968), and south to the mediterranean (madsen 1949). the species is found into the bælt sea (jensen & spärck 1934). occurrence. the boreal and lusitanian regions. habitat. specialised wood-borers. subfossil finds. none. nototeredo norvegica (spengler 1792) distribution. s and w iceland, norway north of lofoten, and south to the mediterranean (madsen 1949). the species extends from the north sea (in driftwood) into the bælt sea as far as kiel (jensen & spärck 1934), and was recorded from the limfjord by collin (1884). occurrence. the boreal and lusitanian regions. habitat. in driftwood. subfossil finds. none. psiloteredo megotara (forbes & hanley 1848) distribution. w greenland, spitsbergen, around iceland, norway north of lofoten, and south to the mediterranean (madsen 1949). in danish waters common on the west coast of jylland and recorded from the kattegat, including øresund (jensen & knudsen 1995). occurrence. the arctic, subarctic, boreal and lusitanian regions, although rare in the last region. habitat. only in driftwood (jensen & knudsen 1995). subfossil finds. none. subclass anomalodesmata order pholadomyoida pandora glacialis leach 1819 distribution. w and e greenland and spitsbergen (madsen 1949). occurrence. the arctic and subarctic regions. habitat. from 2 to 205 m deep on mixed bottoms (ockelmann 1958). only subfossil finds. from the vendsyssel region during early/middle and late weichselian (the older yoldia sea and the younger yoldia sea respectively). fig. 88. a: pholas dactylus linnaeus 1758. skagen 4, 20.0–20.5 m b.s., lab. no. 345,93. × 9.6 fragment with umbonal reflection with septa. mguh 25399. b: pholas dactylus linnaeus 1758. geus bulletin no 3.pmd 28-06-2004, 08:4595 96 lyonsia norwegica (gmelin 1791) distribution. s and w iceland, the faeroes, norway off the lofoten islands, and south to the mediterranean. in danish waters recorded from the kattegat. occurrence. the boreal and lusitanian regions. habitat. from 20 to 250 m deep in sand, silty sand and mud bottoms. subfossil finds. recorded from the skagen region, holocene. in the skagen well recorded from the preboreal and boreal. lyonsia arenosa (möller 1842) distribution. w and e greenland, spitsbergen, and norway north of the lofoten islands. main distribution in the arctic and subarctic regions. habitat. from 3 to about 200 m on mixed bottoms (ockelmann 1958). subfossil finds. recorded from the vendsyssel region during the early/middle weichselian and the late weichselian (the older yoldia sea and the younger yoldia sea respectively). cochlodesma praetenue (pulteney 1799) fig. 89a, b distribution. s and w iceland, the faeroes, norway off the lofoten islands, and south to gibraltar (petersen 1968). the species is rare in danish waters and has only once been taken live in the kattegat (jensen & spärck 1934), although shells are found in the øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from the intertidal zone down to 110 m in sand, mud and gravel bottoms. subfossil finds. the north sea and skagen regions, holocene. from the skagen well recorded from the subatlantic. during the eemian recorded from the skagen region. thracia convexa (wood 1815) distribution. the faeroes, norway off the lofoten islands, and south to the mediterranean (madsen 1949). in danish waters recorded from the kattegat (jensen & knudsen 1995). occurrence. the boreal and lusitanian zones. habitat. offshore down to over 800 m in mud and sand bottoms (poppe & goto 1993). in the danish waters taken between 30–80 m (jensen & spärck 1934). subfossil finds. none. thracia phaseolina (lamarck 1818) fig. 90 distribution. s and w iceland, norway off the lofoten islands, and south to the mediterranean (madsen 1949). the species extends from the north sea (petersen 1977) into the kattegat and øresund, although with few records (jensen & knudsen 1995), and it has been refig. 89. a: cochlodesma praetenue (pulteney 1799). skagen 3, 43.19–43.24, core sample-9. × 9.6. view of the inside of a fragmented right valve. mguh 25401. b: cochlodesma praetenue (pulteney 1799). geus collection. læsø rende, denmark. × 4.8. detailed view of the inside of a right valve showing the resilifer. mguh 25402. geus bulletin no 3.pmd 28-06-2004, 08:4596 97 corded from the limfjord (petersen 1986a). occurrence. the boreal and lusitanian regions. habitat. from the low intertidal zone down to 50 m in fine sand, mud or gravel bottoms (poppe & goto 1993). subfossil finds. the limfjord, north sea, vendsyssel and skagen regions, holocene. from the skagen well recorded from the subatlantic. during the eemian found in the bælt sea and north sea regions. thracia gracilis (jeffreys 1865) distribution. recorded from the atlantic (nordsieck 1969). from the danish waters found in the northern kattegat (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. subfossil finds. none. thracia villosiuscula (macgillivray 1827) distribution. s and w iceland, the faeroes, norway off the lofoten islands, and south to the british isles (petersen 1968). according to poppe & goto (1993) also found in the mediterranean. occurrence. the boreal and lusitanian regions. habitat. in the north sea sampled at depths between 20 and 50 m on hard sand (petersen 1977). subfossil finds. recorded from the north sea region during the eemian. cuspidaria cuspidata (olivi 1792) distribution. norway off the lofoten islands, and south to the mediterranean. in the danish waters found in the kattegat (jensen & spärck 1934) and øresund (jensen & knudsen 1995). occurrence. the boreal and lusitanian regions. habitat. from 20 m to 250 m deep in muddy sand and gravel bottoms (poppe & goto 1993). in danish waters at depths between 30 and 60 m (jensen & spärck 1934). subfossil finds. none. cuspidaria obesa (lovén 1846) distribution. e and w greenland, spitsbergen, around iceland, norway north of the lofoten islands, and south to the mediterranean; however, to the south only at depths greater than 400 m (madsen 1949). in danish waters recorded from the skagerrak and kattegat. occurrence. the arctic, subarctic and boreal regions. habitat. at depths from 40 to 2500 m according to madsen (1949), but in danish waters fairly common in the deeper part of the skagerrak. subfossil finds. none. fig. 90. thracia phaseolina (lamarck 1818). skagen 4, 14.0– 14.5 m b.s., lab. no. 339,93. × 9.6. fragment of left valve. mguh 25403. geus bulletin no 3.pmd 28-06-2004, 08:4597 98 the skagen well the skagen well – perspectives the perspectives of the drilling of the skagen well can be seen by describing the sedimental changes observed in the succession of strata penetrated. however, for the first 30 m of skagen well iii, only wash-samples were taken, in order to establish the well for further drilling down to the prequaternary. therefore, the first drilling segment, composed of sand and gravel, was later repeated elsewhere in order to obtain core samples also from the topmost part. this was done in a nearby position – the so-called skagen iv well – and consequently the full record of shell material and sediments can be given from the core samples obtained from the well, representing all the strata met with in the skagen area from the quaternary, the skagen iii well dgu file no. 1.287. the skagen wells iii and iv are considered to represent one well and are therefore listed together. however, also the wash-samples obtained throughout the quaternary are represented, but only as qualitative analyses with the first occurrences of macrofossils – especially the molluscs – indicated (appendix 2). thirty metres below surface (+ 1 m a.s.), the sediment is well-sorted fine sand. for the following 10 m to 40 m b.s., the average grain size falls within the coarse silt fraction which is consistent down to 75 m b.s. as shown from 11 interjacent measurements, appendix 3. from the depth of 75 m b.s. the average diameter falls within that of medium silt, down to the level of 100 m b.s., still well sorted. at the following levels to a depth of 133 m b.s., the material is fine silt and clay. at a depth of 135 m b.s., fine sand with poor sorting occurs, and at 136–137 m b.s., with very poor sorting and quartiles 40% / 90% of 54.099, which shows a diamict material with a content of stones and only allochthonous shell material, in contrast to the superjacent 130 m. at a depth of 179 m, the well produced a fine-grained material of medium silt, moderately sorted, which is close to what was found above the diamict sequence. with some rise in the average diameter to fine sand and a lowering of the sorting, the next remarkable shift happens at a depth of 187 m b.s., where a new diamict sequence is found down to 195 m b.s. here the boundary to the prequaternary deposits is met with, most probably belonging to the lower cretaceous, according to skagen ii, dgu file no. 1, 43 (sorgenfrei & buch 1964). from the above-mentioned strata in the cored sections, the well can be divided into two parts from the point of view of the present investigation on macrofossils. a sandy to gravelly, clayey to silty wellsorted material found in the upper 130 m of the well and in a smaller interval of almost 10 m at a depth of 179 to 187 m b.s. between these two parts, diamict and well-sorted clayey layers without an in situ macrofauna are found. as the main characteristic feature of the 140 m finegrained and well-sorted material, its content of shellbearing marine molluscs is considered. however, also other marine macrofossils have been recognised but not referred to species level, although recorded on a higher taxonomic level in appendix 3, which represents the finds in the skagen well. also the sedimental data are all shown in appendix 3, allowing a direct comparison between the finds of faunal elements and lithology sensu lato. the organic compound is shown with loss on ignition (550°c) and the occurrences of concretions such as pyrite and iron compounds. also allochthonous shell materials are figured. in consequence of the quantitative analyses of molluscs, diversity and number of specimens are given for the marine strata throughout the whole sequence. it is shown that these figures are very greatly according to the different facies met with. on the basis of the dating within the limits of the carbon-14 method (heier-nielsen et al. 1995), appendix 4 and fig. 3, it is seen that the 130 m thick upper sequence of the well is dated to the last 15 000 14c years. this comprises the whole of the holocene with its boreal sea deposits and the late weichselian with the arctic younger yoldia deposits. but also from 140–150 m b.s., measurements on gas compounds of marine origin (t. layer, personal communication 1999) and shell fragments have been dated, giving ages around 36 000 b.p. 14c years. this means that the diamictic sequence occurring between the two marine strata has taken up material which in age is equivalent to the younger part of the older yoldia clay sediments. the deposition of the diamictic geus bulletin no 3.pmd 28-06-2004, 08:4598 99 sediments is referred to the time of the late weichselian ice advance to the main stationary line in jylland. molluscan shells (macoma calcarea) from the younger part of the older yoldia clay (the macoma calcarea zone, sensu petersen in bahnson et al. (1974)), has recently been dated by the ams method to be in 14c years around 32 000 – 33 000 b.p. (aar-1410 and aar-1411). consequently, it is likely that the older marine strata in the lower marine part of the skagen well can be correlated to part of the sequence demonstrated in the skærumhede well (jessen et al. 1910; bahnson et al. 1974) covering the eemian and the main part of the weichselian. regarding the 130 m thick upper sequence, this is from results of the 14c dates related to the time of the younger yoldia sea and the holocene, and as it appears from the dates of the sediments here is for the first time within the danish area found in a continuous marine succession. this can be explained on the basis of the hitherto unsurpassed thickness of late weichselian–holocene marine strata. therefore, while most of the danish area has a continental period in the time span 11 000 – 7500 b.p. in 14c years (petersen 1985b), the skagen area was so low-lying that it was continuously covered by the sea. this is a reflection of a lesser glacial deposition and the glacio-isostatic down-pressing – the latter amounting to up to 200 m, as seen from the amount of isostatic rebound after the waning of the ice cap over northern denmark (petersen 1990). however, when only holocene dates are used, the estimated rebound of 200 m seems to be greater than the present dates allow when also late weichselian dates are used (petersen 1999). so the low stand of sea level during the latest part of theweichselian and early holocene was not to be overtaken by the isostatic uplift at any time in this area. after the final large eustatic rise in the early atlantic, the marine sedimentation is a dominating factor in raising the levelled sea bottom, compared to a decreasing isostatic rebound up to the subatlantic when the isostatic rebound in denmark expired (petersen 1990). here the formation of the skagen spit takes over, so that the last event of changing depth depends on the large quantities of sand and gravel deposited as the skagen spit grows to the north. the pre-late quaternary deposits in the deeper part of the skagen well, the base of the quaternary is found resting on pre-quaternary deposits of lower cretaceous sand. in the following description, appendix 3 should be consulted. the pre-quaternary strata consist of quartz sand and gravel. from a depth of 195.15–195.30 m b.s., which is the upper part of the pre-quaternary stratum, a mean grain size of fine sand, poorly sorted, is found (lab. no. 789.93). at a depth of 194.35–194.48 m b.s., the sediment is poorly sorted and the mean grain size is within medium sand, and this sediment also contains much quartz, but has another component in the form of stones of granitic composition. the cumulative curve shows an even distribution of all grain sizes, which refers this sediment to be a till (lab. no. 788.93). this is also true for the overlying strata up to around 189 m b.s. the mean grain size is here within fine sand; however, more fine-grained parts are found. the sediment contains siderite(?) concretions with pyrite, in which imprints of cyrena sp. are found. this might show that jurassic deposits have been eroded. some traces in pyrite were found as well. the whole sequence can be regarded as till. the sediment analyses from 188.57 to 187.18 give a badly sorted sediment sustaining that this is a till. also this level contains shell fragments, one of which can be shown to be a nuculana pernula. the late pleistocene eemian deposits the granulometric composition is shown from sample lab. no. 800.93. at a depth of 185.37 m b.s. the sediment is very fine-grained but contains only fragmented bivalves. at 185.0 m b.s. the marine sediment can be demonstrated by the occurrences of dentalium vulgare in many specimens and the bivalve kelliella miliaris also in many specimens and with connected valves. the granulometric composition can be seen from the two levels 182.65 and 180.57 m b.s., samples nos 784.93 and 797.93 respectively. it appears that the sediment is very fine-grained clay to fine silt and moderately sorted in the 180.57 m level. accessory finds of spatangids and ophiorids occur at the 185.0 m level, and pyrite formed in former burrows in the clayey material. also finds of fish occur, as found at the 182 m level, geus bulletin no 3.pmd 28-06-2004, 08:4599 100 and under the name of other fossils also crustacean remains have been listed. the third mollusc species found at this level is the ophistobranch limacina retroversa, which is found in large numbers (11 specimens in one sample) together with kelliella miliaris (also of a number of 15 in one sample). the samples here referred to are all from the heavy weight separation of the foraminifer samples. the species diversity and number of specimens in the sediment appear from the sample at 182.24 m b.s. in which 25 specimens of kelliella miliaris are found – most of them with connected valves and in some parts kept in pyrite. trace fossils in pyrite are found in great quantities recorded in the table from all levels. an expression of the grade of fine-grained sediment occurring at this level can be seen from the fact that only a biogene residue occurs here including the pyritiferous biogene traces – lebenspuren. the limpid delectopecten vitreus also appears at this level. as mentioned in the chapter on the molluscan species, the two sedentary species which today are known from the deeper part of the skagerrak are delectopecten vitreus and kelliella miliaris. the latter is also found in the turritella terebra zone in the skærumhede well. from 183.4 m b.s. hiatella arctica is found, which occurs also at the greater depths and furthermore is a species widely extended. the occurrence of entalina tetragonia at 183.6 m b.s. goes together with the occurrence of delectopecten vitreus, both of which are found in the deeper part of the skagerrak today, where they are part of the amphilopsis norwegica/delectopecten vitreus community. to this can be added cadulus jeffreysi found at the 184.4 m level. this species is widely extended in the northern part of the atlantic down to the bay of biscay and into the mediterranean. however, a single find of siphonodentalium lobatum at 184.6 m b.s. points to a more arctic environment. such shells are found in glaciation deposits according to muus (1959). however, the species may extend into the lusitanian region. from the 182 m level and up to 180 m b.s. still with a fine-grained and well-sorted sediment, yoldiella frigida appears, which is also known from the deeper part of the skagerrak today. this species can be referred to the same environment as has been mentioned above – the amphilopsis norwegica/delectopecten vitreus community. yoldiella frigida is known from the turritella terebra zone in the skærumhede sequence and the portlandia arctica zone according to jensen & spärck (1934). kelliella miliaris and limacina retroversa, which have been very frequent in marine layers met with under 180 m b.s., are no longer found above 180 m b.s. the early/middle weichselian, marine and glacigene deposits regarding the sediment which is to follow at the levels above, between 179.65 and 179.74, it appears that the mean grain size is somewhat bigger medium sand, moderately sorted. but the most significant is found in the cumulative curve showing two maxima on the frequency curve (fig. 91). this points to the effects of two sedimentation agents which might be a drop till effect besides the general marine sedimentation. during the examination of the samples from this level, sand and fine gravel occur, in contrast to the levels below, where only biogene remains were found, including the pyritiferous biogeneous traces. the coarser minerogene elements are found higher up in the series to the level 175.30–175.50 m where a fine-grained sediment with a median grain size of 0.002 mm reveals arctic marine molluscs. this is the first appearance of portlandia arctica, which as the name tells is the characteristic mollusc of the portlandia arctica zone in the skærumhede sequence. however, also the presence of yoldia hyperborea, which is known today from the arctic and down to the lofoten area is characteristic. this species is also found in the portlandia arctica zone together with nuculana pernula and palliolum greenlandicum. the occurrence of macoma sp. has been added from cumulated weight per cent frequency per cent grain size distribution older yoldia sea sediment sample id: 179.65 – 179.74 m 100 90 80 70 60 50 40 30 20 10 0 0. 00 20 0. 00 28 0. 00 39 0. 00 55 0. 00 78 0. 01 10 0. 01 56 0. 02 21 0. 03 12 0. 04 42 0. 09 00 0. 12 50 0. 18 00 0. 25 00 0. 35 50 0. 50 00 0. 71 00 1. 00 00 0. 06 30 w ei gh t pe r ce nt , % grain size, mm fig. 91. the granulometric composition with two maxima on the frequency curve (lab. no. 483.93) at the 179.65–179.74m level. geus bulletin no 3.pmd 28-06-2004, 08:45100 101 the 177.8 m level but not on species level because of the fragmentary state of the shell. both spatangids and ophiorids are found and a fragment of cirriped at the 177.8 m level. at 174.4 m b.s. a single find of yoldia hyperborea occurs. the granulometric composition found at the 173.67–173.85 level (lab. no. 781.93) shows bad sorting in a clayey sediment with a median grain size within clay to fine silt. in this sediment fragments of arctica islandica are found that can be regarded as being part of the redeposited material which can be found also higher up in the core. at a level of 166.5 m b.s. the sediment is well-sorted fine sand and regarded as fluvial. in the following 16 m up to 151.50 m b.s. the sediment is coarser, being a moderately sorted medium sand also regarded as fluvial sand. from here only some shell fragments are found and no record of fossilia varia (other fossils in appendix 3). in the next metres to the level of 143.83–144.00 m b.s. the mean grain size is within the medium silt grade. this is found to be a fine-grained fluvial material forming part of a glacigene complex. also here, fragmentary molluscs are found. the late weichselian marine and glacigene deposits the first molluscs regarded as autochthonous above the glacigene complex are found at the 130.2 m level, so this is regarded as the upper limit of the glacigene sequence. in the interval from 141.00 and up to 130.2 m b.s. more shell fragments have been found – all showing signs of transport and wear. finds of pyrite (137.8 m level), concretion (132.6 m level), and glacial stria on a stone (137.44 m level) all reflect typical features for a till deposit. the marine shell material taken up by the glacier occurs in a fragmentary state, which is typical for redeposited material. however, it is from these strata that the absorption of gases from marine deposits has been dated. these dates form as mentioned a parallel to the age determination of the shells (macoma calcarea) from the skærumhede sequence where the older yoldia clay fauna has been studied (madsen et al. 1908; bahnson et al. 1974). the ages found on macoma calcarea shells from the skærumhede ii well give for the first time, on the basis of molluscs, the absolute age of around 32 000 – 33 000 (14c years) of the younger part of the older yoldia clay. compared with the dating of the marine gases from the skagen well, there is a good correlation to the stratigraphically now well-established skærumhede sequence, so that the two cored sections found on skagen and at skærumhede can be regarded as deposited during the same time in the weichselian. the skagen sequence, however, has been strongly eroded by the ice sheet advancing during the late weichselian. however, the thick packet of up to 50 m glacial sediments consequently contains the traces (gases) of that marine environment, which has been eroded, but is hereby dated to give the maximum age of the glaciation. this age points to the glaciation event in the late weichselian around 20 000 – 18 000 b.p. (petersen & kronborg 1991). however, here the upper marine sequence found in the skagen well will be described. from the 131.63–131.73 m level and up the core the sediment is extremely fine-grained with a medium grain size of fine silt which stays as such a size up to 100 m b.s. it should be noted that throughout the first 15 m of the core from the above-mentioned level finds of coarser material occur. this is seen at the 125.89–126.00 m level (lab. no. 526.93), where the granulometric composition reflects two maxima on the frequency curve (fig. 92). this is taken as a typical sign of a supplementary sedimentation which might have been caused by the melting of floating ice with the coarser sediment imbedded – a drop till effect, as found deeper in the core (fig. 91). this suspected arctic influence is sustained by the occurrences of arctic molluscs up to the 114.0–115 m level, where both portlandia arctica and bathyarca glacialis are present. cumulated weight per cent frequency per cent grain size distribution younger yoldia sea sediment sample id: 125.89 – 126.00 m 100 90 80 70 60 50 40 30 20 10 0 0. 00 20 0. 00 28 0. 00 39 0. 00 55 0. 00 78 0. 01 10 0. 01 56 0. 02 21 0. 03 12 0. 04 42 0. 09 00 0. 12 50 0. 18 00 0. 25 00 0. 35 50 0. 50 00 0. 71 00 1. 00 00 0. 06 30 w ei gh t pe r ce nt , % grain size, mm fig. 92. the granulometric composition with two maxima on the frequency curve (lab. no. 526.93) at the 125.9–126.0 m level. geus bulletin no 3.pmd 28-06-2004, 08:45101 102 furthermore, species such as nuculana pernula, nuculana minuta and yoldiella lenticula occur, which are known from the older yoldia clay in the skærumhede sequence. yoldiella frigida is the first to occur at the 130.2 m level in the skagen well. from this level several finds of ophiuroids (fragments), cirripeds and remains of pisces. however, no finds of spatangoids have been demonstrated within the whole sequence referred to the arctic marine deposits, but they are found all the way up in the boreal sequence (fig. 93e; fold-out, back cover). in the upper part of the arctic sequence siphonodentalium lobatum occurs at the 116.0–114.6 m level and a single find of entalina tetragona. occurrences of nucula sp. and macoma sp. are also recorded in the arctic part, but in such a fragmented state that the species cannot be given. from the recorded faunal composition it appears that it is a deeper-water fauna. this is also supported by the fact that forms reflecting an arctic macoma calcarea community are not present, and the fine-grained sediment points in the same direction. as a comment to the sedimentary environment it should be mentioned that magnetic spherical concretions have been found all through the arctic sequence. from five levels: 117.69–117.85, 124.34–124.50, 127.39– 127.50, 128.13–128.33 and 132.69–132.77 m b.s. a high content of griegite (fe 3 s 4 ), which explains their magnetic quality, has been found by x-ray analysis together with quartz, calcite, feldspar and clay. griegite has been reported as a constituent of reduced sediments. the occurrences in the skagen well are of interest in so far as the spherical magnetic concretions have been recorded only from the arctic sequence. this arctic sea deposit has been dated on material from the cores both foraminifers and macrofossils (heier-nielsen et al. 1995). from this it is seen that the actual time span ranges over 5000 14c years from 15 000 to 10 000 b.p. the sudden change in the macrofauna, or better the abrupt stop of the occurrences of arctic species, at the level of 114.2 to 114.00 m b.s. gives the pleistocene– holocene boundary. the transition from the pleistocene younger yoldia sea to the holocene marine deposits is here recorded for the first time within the danish realm with a whole series of ams datings supporting the chronostratigraphic position, see appendix 4. the dates are highly significant because the molluscan finds in the older part of the marine holocene are extremely poor. this, however, is not caused by the lack of samples from this core section, but is as will be shown dependent on the type of facies following the deposition of the youngest yoldia sea, which was a deeper-water deposit, followed by a boreal deeperwater facies in the older part of the holocene. the change from arctic to boreal conditions is regarded as influenced by a new current system from the atlantic bringing in the new temperate fauna replacing the arctic fauna of late weichselian age. the change in fauna is, however, not reflected in the sedimentary record (appendix 3, pp. 17, 21), which shows a very homogeneous clayey grain size distribution with nothing coarser than fine sand. only in one sample (appendix 3, p. 21, 115 m b.s.) at the sharp boundary between late weichselian and holocene medium sand, coarse sand and gravel are observed. on this homogeneous sequence of clay to fine sand measurements of magnetic susceptibility and thermoluminescence sensitivity have been conducted. it is worth noticing that in a diagram of magnetic susceptibility versus tl sensitivity the two samples forming the peak in the last part of the late weichselian also represent the more immature sediment (high susceptibility and high tl sensitivity). in contrast, the whole series of samples from the lower part of the holocene seems more mature (low susceptibility and low tl sensitivity). so, in this way the peak can also be connected with the sudden break through of the water from the baltic ice lake at mt. billingen, whereas the mature sediments from the holocene may reflect the long-transported sediments introduced by the new current system from the atlantic, bringing in the new temperate fauna in the early part of the holocene and replacing the arctic fauna of late weichselian age (unpublished data, k.l. rasmussen and k.s. petersen). the holocene as mentioned earlier, the transition from the arctic younger yoldia sea to the oldest holocene marine deposits is not to be seen from the sediment analyses except for the occurrences of griegite and some coarser material in the arctic part. this appears when the cumulative curves from the 125.89–126.0 and 113.60– 113.70 m levels from the arctic and boreal part (figs 92, 94, lab. nos 526.93 and 522.93 respectively) are compared. the median grain size is for both samples fine silt, see appendix 3, p. 21. considering the many samples analysed within the geus bulletin no 3.pmd 28-06-2004, 08:45102 103 lower part of the holocene up to the 100 m level, which is dated to be around the boreal–atlantic boundary, only very few molluscan species have been found; also the number of specimens is low. the preboreal–boreal 10 000 – 8000 14c years b.p. in the preboreal–boreal sequence only parvicardium minimum has been found in more than a single find together with mysella bidentata. however, three other genera are recorded: cardium, abra and lyonsia. parvicardium minimum is known from the deeper part of the skagerrak today and is found up to a depth of 30 m in the kattegat. it is recorded also from the eemian in the skærumhede series. compared with the occurrences of mysella bidentata also in this core level at skagen one can imagine a deeper-water environment, because mysella bidentata is also found to great depth (600 m) today in the skagerrak. spatangoids, apparently in great quantities – considering the many fragments – are found and in a lesser degree fragments of ophiuroids, which were also recorded from the arctic part. from the family spatangidae, five genera are known in nordic waters. from the skagen well at a depth of 108.34–108.56 m b.s. a well-preserved species of brissopsis lyrifera (forbes) has been collected (fig. 95). this species lives only on pure muddy bottoms and totally embedded in the sediment. as seen from the grain-size distribution from the level of 107.90–108.00 m b.s., this part is a fine-grained sediment. from the 109.39–109.50 m level the core section revealed a cut through the traces of a spatangoid similar to those that brissopsis lyrifera could leave, with the typical backfilling (bromley 1990, fig. 5.11; see fig. 96). brissopsis lyrifera can be found in great quantities in the northern part of the kattegat and skagerrak, while it might be found in the øresund but not in the bælt sea, the baltic and the limfjord region, according to mortensen (1924). cumulated weight per cent frequency per cent grain size distribution early holocene sediment sample id: 113.60 – 113.70 m 100 90 80 70 60 50 40 30 20 10 0 0. 00 20 0. 00 28 0. 00 39 0. 00 55 0. 00 78 0. 01 10 0. 01 56 0. 02 21 0. 03 12 0. 04 42 0. 09 00 0. 12 50 0. 18 00 0. 25 00 0. 35 50 0. 50 00 0. 71 00 1. 00 00 0. 06 30 w ei gh t pe r ce nt , % grain size, mm fig. 94. the cumulative curve from the 113.60–113.70 m level, lab. no. 522.93. fig. 95. the well-preserved brissopsis lyrifera (forbes) from the 108.34–108.56 m level. mguh 25404. fig. 96. trace from the 109.39–109.50 m level. might be similar to that of brissopsis lyrifera. geus bulletin no 3.pmd 28-06-2004, 08:45103 104 also other spatangoids might be found in the skagen cores in the huge material of fragments. from the older strata the genus echinocardium has been recorded earlier from the skærumhede series by the author, and echinocardium cordatum has been found in the cyprina clay from the eemian (madsen et al. 1908). so poor in molluscan species this community from the early holocene appears to be, one may pay attention to the abundant of remains of starfishes and echinoids which can be seen as a dominating element in this environment. in this way the sea bottom of those days was controlled by the echinoderms eating up most of the larvae of molluscs, as described by thorson (1961). if one should be compared with a present-day community, it must be the maldane-ophiura sarsi community in which besides ophiura sarsi, brissopsis lyrifera is found as the only often found larger animal (thorson 1968). the maldane-ophiura sarsi community replaces the amphiura community at depths of around 150 m and deeper in the skagerrak. a single find of pisces (100.3–100.5 m) has been recorded and a few finds of plant remains and pyritified traces (chondrites?). these rare finds of marine deeper-water facies from the very last part of the pleistocene and the earliest holocene will contribute to our knowledge of the land and sea configuration during the so-called continental period (petersen 1985b). considering the sedimentation rate during the first 2000 years of the holocene, viz.: through the preboreal and the boreal from which there have been only a few records earlier within the danish area, it is seen to be around 7.5 m per 1000 years. this is higher than the sedimentation rate for the younger yoldia sea, as found also in the skagen well record of 3 m per 1000 years. when this is given in calendar years, the differences are even bigger, because then the sedimentation of 15 m in the younger yoldia sea took about 6000 calendar years and still about 2000 calendar years in the preboreal and boreal seas within the skagen area (petersen & rasmussen 1995a, b). regarding the sediment, 50% is found to be clay in the younger yoldia sea – and in some parts at the level of 117.29–117.40 m b.s. around 63% – while during the preboreal–boreal the clay content has fallen from around 40% at the 113.60–113.70 m level to 20– 30% at the 100 m level. the atlantic 8000–5000 14c years b.p. from the dates (heier-nielsen et al. 1995) the atlantic covers the cored section from 100 to 80 m b.s. here the sediment in the oldest part has 30–20% clay, falling to a content of 15% clay in the youngest part at the 80.60–80.70 m level (see fig. 93a). throughout the atlantic the echinoids still dominate the samples and among these the spartangoids, as in the preboreal and boreal. however, here small gastropodsoccur:melanella lubrica,odostomia umbilicaris, and eulimella scillae. melanella lubrica is regarded as an ectoparasite on holothuroids, and odostomia umbilicaris is often found together with mytilus adriaticus. however, the latter bivalve has not been found in the skagen well material. it should be mentioned that the odostomia species are difficult to determine (fretter et al. 1986, p. 605) and no less so in subfossil material. furthermore,onoba vitrea and aclis minor are found in the younger part of the atlantic, where the determination of onoba vitrea is taken with some reservation because the three other species odostomia semicostata, odostomia aculeus, and odostomia proxima are very much alike and difficult to tell apart on shell features alone. aclis minor belongs to a large group of predatory gastropods that mostly and perhaps always (cf. fretter & graham 1962) are associated with echinoderms. from the atlantic single finds of parvicardium minimum from the 96 m level and spisula subtruncata at the 86 m level occur. spisula subtruncata is found next at the 73 m level in the subboreal, but becomes the dominating bivalve at the 30 m level, which can be referred to the younger part of the subatlantic. this depth is also within the range where this bivalve is found in large amount in the present-day danish seas. from the atlantic the predatory gastropod lunatia alderi occurs. this species is most probably the one which has bored into the many molluscs found in the overlying strata, but has not been recognised by its traces in the sparse material from the atlantic. fragments of abra sp. and macoma sp. occur in the cored section from the atlantic, and from the 80.60– 80.70 m level also finds of pisces and crustaceans have been recorded, as seen in appendix 3. it appears that also in the atlantic the sampling reveals a deposit with low diversity and few specimens of molluscs, where the echinoderms dominate as in the preboreal and the boreal sequence. however, considering the older holocene deposits which were tengeus bulletin no 3.pmd 28-06-2004, 08:45104 105 tatively referred to the maldane-ophiura sarsi community, the one from the atlantic can on the basis of the molluscs and the still dominating echinoderms be regarded as another of the deeper-water communities found in the present-day deeper water in skagerrak. here it should be the amphiura community, which as mentioned earlier is found above depths of 150 m. the subboreal 5000–2500 14c years b.p. the following 20 m of the skagen well cover the subboreal, 80 to 60 m b.s. the sedimentation rate can be estimated to be 8 m per 1000 years, a slight rise from the 6.6 m per 1000 years found during the atlantic. the clay content falls in this part to below 10%, and the coarse silt and fine sand fractions become the dominating grain sizes. thematerial iswell sorted. fig. 93a,c. in all the sampled cores within this section fragments of echinoderms occur – mostly spatangoids as found earlier – but the diversity of mollusc species is higher, up to 10 different species in one sample and several with five to seven species in each. however, the number of specimens is still low and most of the finds are of single specimens. only onoba vitrea is found in a number of eight specimens in one sample (the 67.0 m level). among the other species, only lunatia alderi can be mentioned occurring in a number of eight within the whole section. from the 73.0 m level turritella communis occurs with boring of predatory gastropods – probably lunatia alderi. furthermore, eulimella scillae and retusa truncatula are found from the 67 m level and mysella bidentata together with corbula gibba at the 65.6 m level. the latter will be more common in the above-lying strata belonging to the subatlantic. from the subboreal sequence, one of the very few finds of polyplacophora occurs sitting in the sediment, represented, however, only by one plate which does not allow further determination by the author. within the interval from 75.0 m to 72.0 m three finds of turritella communis have been recorded. this is one of the characteristic species on the level muddy bottoms. nuculana minuta which has been found also in the arctic younger yoldia clay is here recorded for the first time in the holocene in the skagen well. there are several finds of nuculana minuta from the subboreal, and it is found in the present-day kattegat on muddy bottoms at depths of more than 20 m. this fits very well with the occurrences of turritella communis. acanthocardia echinata is also found for the first time and here recorded from 78.0 m. this species occurs on mixed bottoms and clay bottoms at depths of from 10 to 150 m. also phaxas pellucidus occurs for the first time in the skagen well during the subboreal. this mollusc occurs in general at depths of between 10 and 50 m, often together with abra alba, also found in this section of the well. the first occurrences of chamelea striatula and corbula gibba are in the skagen well during the subboreal. chamelea striatula is one of the most common of the danish marine bivalves but is connected to the sandy bottoms. according to jensen & spärck (1934), it is not found in the kattegat at depths greater than 50 m, because the sandy bottom in this region goes no further out and the species is rarely taken on clayey bottoms. in this connection it should be noticed that just around the 75 m level, where chamelea striatula occurs for the first time in the skagen well, the sediment changes to coarse silt with more than 50% fine sand. finally, at 61.09–61.14 m, is the first occurrence of tellimya ferruginosa. this species will also be more common in the subatlantic from the 30 m level. tellimya ferruginosa is often connected with the occurrence of echinocardium cordatum but can also be found free living (jensen & spärck 1934). the many new species – new through time in the skagen well – introduced in the subboreal point to water depths around 50 m with characteristic species from the present-day community such as turritella communis and chamelea striatula – the venus community. the changes to a more sandy sediment are perhaps the background for the occurrences of the new species. however, the echinoids have also decreased – and this may explain the more prolific mollusc faunas, for the toll of eaten molluscan larvae taken by the echinoderms is no longer so high (cf. thorson 1961). the subatlantic 2500– 14c years b.p. the uppermost 60 m of the skagen well belongs to the subatlantic. in general the 60 m cored section that falls within the subatlantic can be divided into two parts of an equal length of 30 m: the lower 30 m with geus bulletin no 3.pmd 28-06-2004, 08:45105 106 a clay content of 30–40% of well-sorted sediment, and the upper 30 m mainly consisting of fine to medium sand with few intercalations of gravel. regarding the dated part of this upper sequence (heier-nielsen et al. 1995, table 1) – from 30.25 to 12.75 m b.s. the sedimentation is 17.50 m during 210 years from a.d. 950 to a.d. 1160. this gives a sedimentation rate of about 80 m per 1000 years. in this case showing a fine example of the building up of the skagen spit, where the coarser material occurs as part of the long-shore transport, and deposited in foreset beds. the older subatlantic the older part of the subatlantic covers the interval from 60 to 30 m b.s. this section shows a slight coarsening upwards and a sedimentation rate of 30 m per 1000 years. the faunal composition can be analysed on the basis of 50 samples with a higher species diversity than found in the subboreal. some of the species are new to the record from the skagen well. hinia pygmaea appears for the first time at the 42 m level, as well as hinia reticulata. they both belong to the sublittoral zone and are found on muddy bottoms. mangelia brachystoma, which first occurred at the 58 m level, also belongs to the sublittoral fauna, but it occurs on sand and sandy muddy bottoms. polygireulima sinuosa is an ectoparasite on echinoderms which are still common and constitute a part of every one of the samples, but it has been found only within the level 38.19–38.24 m. the first littoral species, mytilus edulis, occurs at 49.14–49.24 m, and from this level it occurs regularly upwards, but only in small numbers until the 31 m level, where it is found in greater quantities. this species can be found out to 40 m depth, but must nevertheless be considered an eulittoral species where its occurrence is most abundant. the young specimens are often found on the vegetation. also chlamys varia is common in the coastal zone and occurs at the 32.85–32.90 m level. heteranomia squamula is epifaunal on hard substrates but also on algae and crustaceans. it has a wide occurrence from the littoral zone out to a depth of 100 m. in the skagen well it is confined to the subatlantic part. thyasira flexuosa is today a common bivalve on clayey bottoms from 20 m to 100 m, but it has been found only in two samples from the older subatlantic. this is hard to explain, as it has a wide extension within the whole of the north atlantic area (jensen & spärck 1934), and in numbers it is one of the most dominant species on the muddy bottoms which according to the sediment analysis have been prevalent for most of the holocene in the skagen area. turtonia minuta, belonging to the species from the coastal zone, is found in a single specimen at 39.85– 40.02 m. it is not recorded from the recent danish fauna, but lives off the norwegian west coast and is found subfossil in the limfjord region. at 47.30–47.35 m is the youngest record of parvicardium minimum, which was one of the few species occurring in the older holocene reflecting deeper water. a single find of angulus tenuis is at 55.30–55.35 m. the common occurrence of this species starts at the 30 m level. also donax vittatus occurs at 35.90–36.00 m level which must be seen as outside the general occurrence of this species, which is from littoral to around the 20 m depth donax vittatus is found within the high energy zone. a single find of abra prismatica at 49.90–42.00 m is within the general depth interval for this species (20– 60 m). in connection with the depth indications given in the well in metres below surface and the common depth intervals indicated by various authors for the molluscan species, it is possible to use the actual depth recorded in the well as the living depth for the subfossil fauna found in the skagen well during the younger part of the holocene. this because of the expiring isostatic movement and only little eustatic changes during the late holocene (petersen 1991b). corbula gibba, which was also found during the subboreal has in the subatlantic an even occurrence through the older part. barnea candida, normally only found out to a depth of 30 m, occurs in the well already at the 51.54–51.59 m level, although only found in fragments. cochlodesma praetenue which was found in the eemian at 183.77–184.00 m b.s. is also found in the subatlantic at 43.19–43.24 m. this species is rare in danish waters and has been taken alive only once north-east of the island of læsø. however, shells have been found elsewhere in the kattegat region, jensen & spärck (1934). it has a wide occurrence from the littoral zone and out to 110 m on different bottom types. finally thracia phaseolina shall be mentioned. this species occurs to depths around 50 m on clayey bottoms. as mentioned above, the echinoderms are also found geus bulletin no 3.pmd 28-06-2004, 08:45106 107 in the subatlantic represented by the fragments of spatangoids. also cirripeds occur in still higher quantities up towards the 30 m level (appendix 3, p. 6). furthermore, there are single finds of pisces and other fossil remains such as crustaceans (other fossils in appendix 3). however, also serpulae are found and may have settled on the shells of the other animals as the crustacean carapax. the younger subatlantic the increasing number of cirripeds in the upper 30 m should probably be regarded as allochthonous, since they occur together with the coarser material during the formation of the advancing skagen spit. the change in the upper 30 m to coarser material also introduces new forms of molluscs that are characteristic of the littoral facies and high-energy coastal situation still prevailing in this area today. the description of the upper 30 m is, as mentioned earlier, based on the skagen iv well 50 m away from the skagen iii well and at the same level (+ 1 m). this was done because only washed samples were obtained from the upper 30 m of the skagen iii well, and such samples could not form the best basis for a uniform description of the whole sequence – especially the necessary quantitative treatment of the molluscan faunas could not be fulfilled in that way. furthermore, a total of 29 grain-size analyses have been made within this part of the column, showing two sequences of well-sorted sediment coarsening upwards, with a sorting coefficient lower than 2 (fig. 97). in order to control the degree of transported shell material, size analyses and counts on right and left valves have been considered relevant with such a highenergy near shore sedimentation (fig. 98). especially the most prevalent bivalve within these uppermost 30 m, spisula subtruncata, has been counted. also observed borings have been figured in appendix 3, to be seen in connection with the actual finds of the predatory gastropods. this is done in order to show the degree of mutual connection in the molluscan assemblages, between predatory elements and their prey. the building-up of the upper 30 m took place within a very short period of time, and the sedimentation rate of this interval is estimated to be around 70 m per 1000 years. this high sedimentation rate has a serious effect on animal life. thedates on the building-up of the skagen spit lead to the conclusion that the extension of the coast line to the place where the wells have been sunk took place around a.d. 1400. taking into account that the final history of the coastal development takes place as a near-shore and littoral deposition history, the actual development on a west coast site similar to the skagen area has been analysed. this has been done by way of several van veen grab samples – altogether 61 outside the agger tange complex in the westernmost part of the limfjord (petersen 1994a). these investigations focused on the bivalves, evaluating their degree of being autochthonous from the preservation with both valves together, one valve but whole, a fragmented state or a rolled fragment. these observations have been summarised in appendix 5. the newcomers of molluscs from the skagen well will be mentioned. these also represent the species earlier known to live close to the recent danish waters and species new compared to what is known to be part of the recent danish fauna. this part of the record has the highest diversity and number of specimens compared to other sections of the skagen well. the mean species diversity per sample shows a rise compared to the older part and reflects the new sedimentary facies. however, the near to shore situation also puts forward the question of whether part of the faunas, if not all, may have been reworked. eliminating the uppermost ten samples covering the 5 m which can be regarded as the medieval shore. first the species represented by only few finds that are commonly found in great quantities will be discussed. lacuna pallidula occurs only as a single find at 30.0– 30.5 m level. this species occurs on fucus serratus and in great quantities from the littoral and to a depth of 70 m. hydrobia ulvae occurs normally in high numbers in shallow water. in the skagen well it has been recorded from only two levels (11.70–11.80 m and 25.0– 25.5 m) and with few specimens. rissoa violacea is connected with seaweeds and found from the tidal zone to a depth of 50 m. here the only finds are from the 27.0–27.5 and 28.0–28.5 m levels. also bittium reticulatum appears not to be part of the environment, since this species has only one occurrence at the 22.0– 22.5 m level. this species lives on zostera, as do other of the above-mentioned species. it can be concluded that the upper 30 m section of the well lacks the normal abundance of epifaunal elements connected with vegetation. this is also in good accordance with the high rate of sedimentation. geus bulletin no 3.pmd 28-06-2004, 08:45107 108 0 10 20 30 40 50 60 70 80 90 100 weight per cent, % d ep th b el ow s ur fa ce , m la bo ra to ry n um be r clay/silt medium sand gravelfine sand coarse sand 2.2 20.8 3.6 4.6 5.6 6.6 7.6 8.6 9.6 10.6 11.6 12.6 13.6 14.6 15.6 16.6 17.6 18.6 19.6 21.6 22.6 23.6 24.6 25.6 26.6 27.8 28.4 29.8 30.6 295.93 296.93 297.93 298.93 299.93 300.93 301.93 302.93 303.93 304.93 305.93 306.93 307.93 308.93 309.93 310.93 311.93 312.93 313.93 314.93 315.93 316.93 317.93 318.93 319.93 320.93 321.93 323.93 324.93 histogram of 29 grain-size analyses from the upper 30 m of the skagen well 4 among the gastropods occurring in the upper part of the well, aporrhais pespelicani occurs in the interval from 22.0–22.5 m to the 11.0–11.5 m level. this species is regarded as sublittoral from depths of 10– 180 m on a sandy muddy bottom or muddy bottom. however, it has been found in large quantities as empty shells on the shore of the east coast of skagen. this was rather puzzling until it was explained that the hermit crab might have been the actual agent bringing the shells on shore (g.h. petersen, personal communication 1998). the occurrence of lunatia montagui is restricted to the 20.0–20.5 m level, while lunatia alderi is rather frequent in the core samples. the impact of these predatory gastropods on the fauna – 15 species have been recorded with such borings, including some of the lunatia species themselves – has been quantified in appendix 3. the high number of lunatia alderi in the upper 30 m is in accordance with the preferred environment of clean sand of this species. a new neogastropod to the fauna of the well is the buccinum undatum from the 11.0–11.5 m level, while hinia pygmaea now becomes common, occurring in most of the samples from the 28.0–28.5 m level to 6.0– 6.5 m b.s. and represented in many specimens – some of them bored by predatory gastropods, as seen in appendix 3. the small gastropod oenopota turricula has a wide depth range (20–200 m), so the single finds at the 23.5– fig. 97. histogram of 29 grain-size analyses from the upper 30 m of the skagen well 4, showing two coarsening-upwards sequences. geus bulletin no 3.pmd 28-06-2004, 08:45108 109 21.0 m level most probably reflect that only in this part of the well does the clean sandy bottom occur which is preferred by oenopota turricula. of the heterogastropod newcomers in the upper section, graphis albida from the 25.0–25.5 m level can be mentioned. this species is not recorded among the recent danish molluscs (jensen & knudsen 1995). it is found sublittorally out to a 30 m depth. hemiaclis ventrosaoccursat the30.0–30.5mand11.0– 11.5 m levels, but it is recorded in recent waters at a much deeper level: 100–200 m. neither this nor the species mentioned above is recorded from danish waters. vitreolina philippii, occurring within the interval from 29.5 to 7.0 m with seven specimens, is known from the recent danish fauna and is noted as sublittoral to a depth of 200 m. this gastropod is a parasite on echinoderms, as the other eulimidae. echinoderms are still present in the material as seen from appendix 3. from the 15.0–15.5 m level, finds of chrysallida decussata occur, which is also recorded by jensen & knudsen (1995). this species occurs at the depth interval of 14–40 m. turboniella acuta has been recorded from danish waters by jensen & knudsen (1995) although rare. the occurrence of this species in the skagen well is at the 21.0–21.5 m and 20.0–20.5 m levels with, two wellpreserved specimens. among the ophistobranchs there are some fragmentary finds which have not been identified to species ratio 1:1 rightleft > 4.0> 0.5 > 2.0 > 6.5 > 8.0 size, mm 125 48 109 147 311 321 442 569 278 1129 4332 1201 113 85 166 826 98 68 59 14 0 79 75 92 39 5.0–5.5 6.0–6.5 7.0–7.5 8.0–8.5 9.0–9.5 10.0–10.5 11.0–11.5 12.0–12.5 13.0–13.5 14.0–14.5 15.0–15.5 16.0–16.5 17.0–17.5 18.0–18.5 19.0–19.5 20.0–20.5 21.0–21.5 22.0–22.5 23.0–23.5 24.0–24.5 25.0–25.5 26.0–26.5 27.0–27.5 28.0–28.5 29.0–29.5 depth below surface, m number of valves relative size distribution left/right valves ratio valves: spisula subtruncata fig. 98. size histograms for spisula subtruncata in the upper part of the subatlantic sequence, the 29.0–29.5 m level to the 5.0–5.5 m level, with ratio on left and right valves from the 29.0–29.5 m level to the 11.0–11.5 m level. geus bulletin no 3.pmd 28-06-2004, 08:45109 110 level,but species suchasretusa truncatulaandcylichna alba are found also in the upper part of the well. a fragmentary scaphopod from the 15.0–15.5 m level has not been referable to species level. among the bivalves, many are new to the already mentioned fauna from the well, and the number of specimens is for many of the species very high in comparison to what has been recorded from the older strata. of palaeotaxodonta, nucula nitidosa is found and represented all through the interval from 29.5 to 13.0 m b.s., occurring on sand bottom, which is the preferred substrate. also nucula nucleus is found within the interval from 30.5 to 8.0 m b.s. with many (13) specimens, part of them bored as the presiding species by the predatory gastropods. in the subclass pteriomorphia, species from mytiloida and pterioida such as musculus discors at the 28.0– 28.5 level and mytilus edulis in large quantities (113 specimens) are found, albeit most of the latter as juveniles. from the 28.0–29.5 m level individuals are found (with both valves). this latter species is typical in the littoral zone, but may occur at depths out to 40 m. pectinidae have been found, but all in fragments, in the interval 25.5–12.0 m b.s. ostrea edulis occurs in the interval 28.5–7.0 m b.s. – mostly as juveniles. the subclass heterodonta, from where most of the found bivalves come also includes the species most often found and characteristic of the youngest part of the marine sequence. mysella bidentata is recorded from the entire holocene, although only a few specimens are present in the early holocene. in the latest holocene as the present 30 m, 125 specimens have been found. the closely related tellimya ferruginosa occurs apart from a single find at the 61.09–61.14 m level, from the 29.5 m level where it is common up to 8.0 m b.s. both of these species have specimens bored by the predatory gastropods. tellimya ferruginosa is a commensal on echinocardium cordatum, but can also be found on its own in the sediment. mactra stultorum has been found only in the upper part of the cored section and can be seen as connected with the clean sand that is the type of bottom preferred by this species. on a suitable bottom it may be found out to a depth of 60 m. spisula subtruncata, which has a wide distribution from the littoral zone and out to a depth of 200 m, can be found both on muddy and on sandy bottoms. it dominates the uppermost part of the sequence, with 11085 specimens! in recent waters on sandy bottoms this species is one of the most common bivalves in the kattegat at depths between 20 and 30 m (jensen & spärck 1934). on the cored material from the skagen well size histograms and counts on left and right valves have been made in order to ascertain from such measurements whether the shell material is autochthonous/ parautochthonous. as seen from the figures in fig. 98, it appears that there is an even representation of the left and right valves, and the size histograms reflect a life assemblage which also might appear from the wellpreserved gracile valves. the borings counted on valves of this species make it clear that spisula subtruncata must have been the preferred victim of the predatory gastropods in this molluscan assemblage. at the 15.0–15.5 m level around 10% of the specimens are bored (2723 individuals out of which 268 have been bored). individuals (with both valves) have been found up to the 21.0–21.5 m level, where also other bivalves have been found with closed valves. however, the most even occurrence of left– right valves also at the 15.0–15.5 m level (2147–2105 respectively) may speak in favour of an autochthonous state also at this depth. the size histogram from the same level points to the same conclusion (see fig. 98). by way of the same kind of measurements it is possible to extend the possible life-assemblages up to a level of 10.0–10.5 m b.s., where the material still is present in such quantity that the measurements can be taken as bearing. investigations performed on nearshore deposits off the west coast of jylland in the agger tange area given in appendix 4 support the view that lifeassemblages can be found near to shore at depths of up to 6–7 m. almost all the ams datings in the upper part of the well have been based on spisula subtruncata, and these datings all fall within the right relative age according to their respective levels. this is not the case with the date on donax vittatus, which has also been dated within the upper 30 m interval. as shown on the dating diagram (heier-nielsen et al. 1995; appendix 3), the donax vittatus age clearly appears as an older element in a younger part of the section. however, donax vittatus will be commented upon later in the text. solenidae species often occur in the upper part of the sequence, butoften in a fragmentary state.however, phaxas pellucidus is common in the interval between the 30.5 and 20.0 m level, where it is found in several specimens in some of the samples. it lives on different bottom types from the sublittoral at a depth of 4 m out to a depth of 150 m. however, in the skagen well there is only a single occurrence at 73.39–73.44 m b.s. geus bulletin no 3.pmd 28-06-2004, 08:45110 111 one of the dominating bivalves is fabulina fabula, which only occurs within the interval 28.5–4.0 m b.s. some of the specimens have been the victims of the predatory gastropods. this species prefers a bottom type of fine sand, which might explain the interval of occurrence in the skagen well, where there are sandy layers only in the uppermost 30 m. on a suitable bottom this species goes out to a depth of 50 m. also tellina pygmaea and angulus tenuis occur in the upper part of the sequence and only there, with the exception of a single find of angulus tenuis at the 55.30–55.35 m level. this is outside the general occurrence of this shallow-water species normally found from the intertidal zone out to a depth of 10 m. donax vittatus, which is found regularly in the interval from the 27.5–4.0 m level, but often in a fragmentary state, is a typical high-energy coastal form on a sandy bottom. as already mentioned in connection with the dates, donax vittatus also occurs as an allochthonous element, which can be seen from the many rolled fragments of this sturdy shell. its occurrence out to a depth of 20 m off high-energy shores characterises in the best way the situation by the building up of the skagen spit system. the species is not found in the kattegat region and is absent from the inner part of the limfjord. gari fervensis is found only in this upper part of the skagen well from the 27.5 to 23.0 m level. accordingly, in danish waters it is known from a depth of 15–40 m on mixed bottoms and sometimes on sandy bottoms. through most of the holocene, fragments of the genus abra have been found. abra nitida, which has a wide depth distribution from the sublittoral zone out to a depth of 200 m, has been found through the last part of the holocene from the 71.89–72.00 m level to the10.70–10.80m level mostly in single specimens. this species has its maindistribution today in thedeeperparts of the skagerrak and the kattegat on muddy bottoms. a single find of a rolled fragment of arctica islandica occurs at 14.70–14.80 m, which is the only find besides the fragment from the glacial series at the 173.67– 173.85 m level. however, the washed samples have given another specimen also from the subatlantic (appendix 2). chamelea striatula is the characteristic animal of the venus community on a sandy bottom in the north sea and the kattegat. at skagen it occurred already at the 76.34–76.50 m level (late subboreal). at this depth a change of weight per cent of clay takes place (from 13.7% to 7.8%), and the fine sand component becomes the dominating grain size with a weight per cent of 54.9. from the 30 m level, chamelea striatula is more common, and specimens with connected valves occur up to 21.0–21.5 m b.s., many of them bored by predatory gastropods, as shown in appendix 3. from the point of view that also other bivalve species have been found as whole individuals up to the 20 m level, it can be regarded as the well-established limit for an autochthonous occurrence of the molluscs. however, as seen from the observations off the agger tange area given in appendix 5, there will always be an element of allochthonous material in such a high-energy coastal environment which should be taken into account also for the skagen area regarding the uppermost part of the sequence from the skagen well. a single find of timoclea ovata is also found in the upper part of the section at the 27.0–27.5 m level. this species is today found at a greater depth than chamelea striatula, but is not very numerous. within the order myoida, corbula gibba is also well represented in the upper section, with individuals found up to a level of 27.0–27.5 m b.s. this species also shows many specimens killed by predatory gastropods. corbula gibba is found in the sublittoral zone out to a depth of 250 m. at skagen its first occurrence is at 74.89–75.00 m, in the early subboreal, but it becomes common in the subatlantic and occurs in high numbers only in the last part of the subatlantic from the 30 m level, often bored. finally, two single finds of saxicavella jeffreysi and pholas dactylus occurred at the 20.0–20.5 m level. saxicavella jeffreysi is in recent danish waters not very abundant at depths between 25 and 50 m, while pholas dactylus would only be expected to be found at depths less than 10 m. pholas dactylus is a boring form found in hard substrates, which is far from the actual sediment occurring at this level in the skagen well. the fragmentary pholas dactylus can be regarded as one of the allochthonous elements that can be seen in connection with the accessory finds mentioned in appendix 3 and commented upon below. among the accessory finds the barnacles and sea urchins dominate. also fish remains are found, often in the form of vertebrae, but an otolith appears as well. other fossil remains are serpulids, bryozoans, crustaceans, and plant and insect remains, which taken as a whole very well characterise the near-shore environment. on the other hand, no concretions are found like the ones from the younger yoldia sea, or pyrite as found at the base of the holocene and the eemian. although these accessory elements cannot be quantigeus bulletin no 3.pmd 28-06-2004, 08:45111 112 fied, they offer some additional information when considered together with the sedimentological and mollusc records. conclusive remarks on the skagen well in the description given above, the faunal record is the basis for understanding the climatic changes in the skagen well, supplemented by the observation on the changes in the sediments. however, the changes found during the holocene are most likely to be connected with changes in facies, and here the changing depth is the most prominent agent, ending up with the last event represented by the depositional history of the skagen spit. based on the dating of the holocene and the late weichselian, the descriptions have been given in terms of episodes. especially the holocene strata points to a development from deeperto shallow-water facies from preboreal to subatlantic. in this development there appears to be a facies change that can be compared to the bottom communities as known from the skagerrak–kattegat region when going from the deeper-water communities of the present day, the so-called maldane-ophiura sarsi community, to the venus community of the more shallow seas. the mollusc assemblages in the skagen sequence indicate a deeper-water facies during the eemian, the weichselian and the older holocene in contrast to what hitherto was known in other parts of the danish area during the late quaternary. the skagen well has a record of the changing seas during the late quaternary, from the eemian through the weichselian (although only in parts) and the holocene. for the first time within the danish area, the full record of the marine environmental transition from the late pleistocene to the holocene can be demonstrated on the basis of molluscs. however, not all the episodes known from the skagen well can be found in marine facies of the other regions, but thanks to the new records from the north sea around the jydske rev area, a near to full holocene marine record is at hand, including part of the preboreal (petersen 1998). the environmental changes through time in the seven sectors based on the molluscan records the recorded mollusc species within each area are given in appendix 6. regarding the environmental changes through time within the danish realm, the seven sectors will be considered from the eemian, starting in the south within the classical area where forchhammer named the deposits the cyprina clay. eemian species sorted after climatic affinities the bælt sea age: eemian climatic regions: asbl class bivalvia subclass heterodonta order myoida mya truncata linnaeus 1758 total for climatic regions asb. : 1 (1.7%) climatic regions: .sb. class bivalvia subclass heterodonta order veneroida arctica islandica (linnaeus 1767) order myoida zirfaea crispata (linnaeus 1758) total for climatic regions .sb. : 2 (3.4%) climatic regions: .sbl class bivalvia subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 order pterioida geus bulletin no 3.pmd 28-06-2004, 08:45112 113 heteranomia squamula (linnaeus 1758) subclass heterodonta order veneroida mysella bidentata (montagu 1803) tellimya ferruginosa (montagu 1803) spisula elliptica (brown 1827) macoma balthica (linnaeus 1758) total for climatic regions .sbl: 6 (10.2%) climatic regions: ..bl class gastropoda subclass prosobranchia order neotaenioglossa littorina littorea (linnaeus 1758) hydrobia ulvae (pennant 1777) rissoa inconspicua alder 1844 rissoa membranacea (j. adams 1800) rissoa parva (da costa 1779) caecum glabrum (montagu 1803) bittium reticulatum (da costa 1778) aporrhais pespelicani (linnaeus 1758) order heterogastropoda triphora adversa (montagu 1803) epitonium clathrus (linnaeus 1758) order neogastropoda hinia pygmaea (lamarck 1822) hinia reticulata (linnaeus 1758) subclass heterobranchia order heterostropha odostomia scalaris macgillivray 1843 chrysallida obtusa (brown 1827) chrysallida spiralis (montagu 1803) ebala nitidissima (montagu 1803) turbonilla crenata (brown 1827) turbonilla lactea (linné 1758) subclass opisthobranchia order bullomorpha philine aperta (linnaeus 1767) order anaspidea retusa truncatula (bruguière 1792) retusa umbilicata (montagu 1803) akera bullata müller 1776 class bivalvia subclass palaeotaxodonta order nuculoida nucula nitidosa winckworth 1930 nucula sulcata (bronn 1831) subclass pteriomorpha order mytiloida modiolula phaseolina (philippi 1844) order pterioida ostrea edulis linnaeus 1758 subclass heterodonta order veneroida lepton nitidum (turton 1822) acanthocardia echinata (linnaeus 1758) parvicardium exiguum (gmelin 1791) cerastoderma edule (linnaeus 1758) mactra stultorum (linnaeus 1758) spisula subtruncata (da costa 1778) ensis ensis (linnaeus 1758) angulus tenuis (da costa 1778) scrobicularia plana (da costa 1778) abra alba (wood 1802) abra prismatica (montagu 1803) chamelea striatula (da costa 1778) timoclea ovata (pennant 1777) dosinia lincta (montagu 1803) order myoida corbula gibba (olivi 1792) barnea candida (linnaeus 1758) subclass anomalodesmata order pholadomyoida thracia phaseolina (lamarck 1818) total for climatic regions ..bl: 43 (72.9%) climatic regions: ...l class gastropoda subclass opisthobranchia order bullomorpha haminoea navicula (da costa 1778) class bivalvia subclass pteriomorpha order mytiloida mytilaster solidus (poli 1795) subclass heterodonta order veneroida lucinella divaricata (linnaeus 1758) gastrana fragilis (linnaeus 1758) abra segmentum (récluz 1843) paphia senescens (cocconi 1873) gouldia minima (montagu 1803) total for climatic regions ...l: 7 (11.9%) total for the eemian bælt sea: 59 (23.9%) fifty-nine species have been found in the bælt sea region during the eemian, seven of which are found or have been, as is the case with paphia aurea senescens, in the lusitanian region. all of the lusitanian species are found only in the eemian deposits and represent species living in the shallow-water environgeus bulletin no 3.pmd 28-06-2004, 08:45113 114 ment. mytilaster solidus is intertidal attached to rocks or algae, abra segmentum is infralittoral on sandy mud, and haminoea navicula found in zostera beds in sheltered areas. the other three living species lucinella divaricata, gastrana fragilis and gouldia minima are found from or just below the tidal zone and further out at different depths. paphia aurea senescens may as well be regarded as a shallow-water species, considering the distribution of the other tapes species. in this way the overall climatic characterising species of the eemian in the bælt sea area are connected with the shallow-water environment. all the boreo-lusitanian species are known from the recent danish fauna. this is by far the largest group of molluscs, with 43 species forming 72.9% of the eemian bælt sea fauna. 33 species can be found in the tidal to shallow-water environment, while 11 species – epitonium clathrus, ebala nitidissima, turbonilla crenata, retusa umbilicata, akera bullata, nucula nitidosa, nucula sulcata, lepton nitidum, acanthocardia echinata, phaxas pellucidus and timoclea ovata – find their minimum depth, defined by acanthocardia echinata, phaxas pellucidus and timoclea ovata, at 4 m, and epitonium clathrus at 5 m. so rather considering the maximum depth indicated by some of the shallow-water species, there must be two faunas, of which one is shallow out to a depth of a few metres and another for deeper water. there are six species with a rather broad range from the subarctic to the lusitanian regions. five of these species can be found in shallow water, including the intertidal zone, except spisula elliptica, which occurs only at greater depth. the eulittoral species mytilus edulis has given name to the mytilus beds found in the lower part of the marine eemian deposits in the bælt sea region, characterising the littoral deposits. only two species, arctica islandica and zirfaea crispata from the bælt sea region, have a subarctic–boreal distribution. the overall characteristic species for the eemian bælt sea deposits – arctica islandica – can be found from the intertidal zone to great depth, but in the inner recent danish waters often at depths from 10 to 60 m. in the eemian this species characterises the clay deposited during the deeper-water facies. the other mainly boreal species zirfaea crispata has been found only in the upper part of the tapes sand at stensigmose (madsen et al. 1908, p. 176) which fits well with the depth interval of this species from low tide to 7 m. finally mya truncata covers the arctic, subarctic, boreal and lusitanian regions down to the bay of biscay and can be found from the intertidal zone down to 75 m, in danish waters often between 10 and 20 m. taken together, all the information from the abovementioned climatic groups indicates that the bælt sea deposits are represented by two facies. one in littoral/ infralittoral water not deeper than maximum 10 m, and one at depths of more than 5 m, and it is seen that all of the characteristic eemian species within the bælt sea area are lusitanian species connected with the shallow-water facies. fifteen mollusc species of the eemian bælt sea fauna show by the region of lowest mean salinity they inhabit that the salinity of the bælt sea area must have been higher than present-day waters by up to 30–33‰ (sorgenfrei 1958, table 11). here listed as mentioned by nordmann (1928, pp. 79–81): 33‰ circe minima (gouldia minima), 31‰ dosinia lincta, 33‰ lepton nitidum, 33‰ mactra stultorum, 30‰ montacuta ferruginosa (tellimya ferruginosa), 30‰ mytilus phaseolinus (modiolula phaseolina), 33‰ nucula sulcata, 31‰ syndosmya prismatica (abra prismatica), 30‰ venus gallina (chamelea striatula), 30‰ parthenia interstincta (chrysallida spiralis), 32‰ rissoa parva, 30‰ (epitonium clathrus), 33‰ turbonilla lactea, and 30‰ turbonilla rufa (turbonilla crenata). in the recent bottom communities out in the bælt sea area, macoma balthica is the overall characteristic mollusc occurring in all the samples from this region, in some cases being the only mollusc and in some cases together with others such as mytilus edulis, cerastoderma edule and scrobicularia plana in shallow water and in places with a vegetation of gastropods like littorina littorea, littorina tenebrosa, rissoa membranacea and rissoa inconspicua (petersen 1913). the littoral elements with mytilus edulis are well documented in the eemian deposits by the so-called mytilus beds, together with the infaunal paphia aurea senescens found in situ within the strata around the mytilus beds or even down in the freshwater layers. however, macoma balthica is only found in the tapes sands at stensigmose in southern jutland (fig. 1) and not at the many other bælt sea localities (nordmann 1908, 1928). in the deeper-water community of today the macoma community is replaced by the abra alba or astarte communities (petersen 1913). here again the eemian deposits differ on the leading species in that fig. 99. the voderup klint section on the island of ærø with dislocated marine clays of eemian age. geus bulletin no 3.pmd 28-06-2004, 08:45114 115 geus bulletin no 3.pmd 28-06-2004, 08:45115 116 fig. 100. one of the cyprina clay outcrops along the voderup klint profile on the island of ærø. abra alba is recorded only from stensigmose and the astartidae tridonta borealis, tridonta elliptica and tridonta montagui have not been found at all in the bælt sea region or further to the east within the baltic. the characteristic species of the eemian deeperwater deposits, the cyprina clay (with an abundant number of arctica islandica), is also present in the recent deeper water of the bælt sea, but petersen (1913, p. 4) avoided using this bivalve as one of his characterising species in his ‘evaluation of the sea’ because of its uneven distribution, which is certainly not the case considering the present outcrops of the cyprina clay along the shores of the bælt sea region (figs 99, 100). the great difference between the eemian bælt sea fauna and the recent one which has been outlined above can be considered together with the concluding remarks of nordmann (1908, pp. 113 and 148) on cyprina islandica, stating that this species cannot be regarded as a characterising fossil species of the eemian deposits, but more probably should be seen as a relict in the inner part of the eemian waters from the sea, of a more arctic–boreal nature, so to say forming a parallel to the occurrences of the astarte species in the present bælt sea and baltic regions. the baltic age: eemian climatic regions: asb. class bivalvia subclass heterodonta order veneroida macoma calcarea (gmelin 1791) total for climatic regions asb. : 1 (5.3%) climatic regions: asbl class bivalvia subclass heterodonta order myoida mya truncata linnaeus 1758 total for climatic regions asbl: 1 (5.3%) climatic regions: .sb. class bivalvia subclass heterodonta order veneroida arctica islandica (linnaeus 1767) total for climatic regions .sb. : 1 (5.3%) climatic regions: .sbl class bivalvia subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 modiolus modiolus (linnaeus 1758) subclass heterodonta order veneroida mysella bidentata (montagu 1803) macoma balthica (linnaeus 1758) total for climatic regions .sbl: 4 (21.1%) climatic regions: ..bl class gastropoda subclass prosobranchia order neotaenioglossa littorina littorea (linnaeus 1758) hydrobia ulvae (pennant 1777) turritella communis risso 1826 aporrhais pespelicani (linnaeus 1758) lunatia alderi (forbes 1838) order neogastropoda hinia reticulata (linnaeus 1758) class bivalvia subclass palaeotaxodonta order nuculoida nucula nitidosa winckworth 1930 nucula nucleus (linnaeus 1767) subclass heterodonta order veneroida cerastoderma edule (linnaeus 1758) spisula subtruncata (da costa 1778) scrobicularia plana (da costa 1778) order myoida corbula gibba (olivi 1792) total for climatic regions ..bl: 12 (63.2%) total for the eemian baltic: 19 (7.7%) the baltic sea fauna comprises only 19 species. however, five are new here compared to the eemian known from the bælt sea area: three from the boreo-lusitanian climatic region, i.e. turritella communis, lunatia alderi and nucula nucleus; one from the subarctic– lusitanian region, i.e. modiolus modiolus; and one with an arctic–boreal distribution, i.e. macoma calcarea. according to the living depth of turritella communis geus bulletin no 3.pmd 28-06-2004, 08:45116 117 geus bulletin no 3.pmd 28-06-2004, 08:45117 118 and lunatia alderi, this should be an indication of environment deeper than 10 m. but the eulittoral mytilus edulis and cerastoderma edule are nevertheless recorded from the baltic sea area, although these species do not come from layers in which they occur in great quantities in the bælt sea area (the mytilus beds). from the climatic affinities of the molluscan finds in the baltic sea area, it appears that no pure lusitanian species are found. this is the biggest contrast to the bælt sea region. it has also been the basis for keeping these finds apart, as done by ødum (1933). however, considering the new element in this fauna, turritella communis, compared to the bælt sea fauna, it has been argued by way of correlation on the basis of the foraminifera that these deposits can be regarded as eemian and that the facies belongs in deeper water than known from the bælt sea area (petersen & konradi 1974). such deposits occur, besides that at strandegaard dyrehave, also at some places on møn characterised by the occurrences of turritella communis (berthelsen et al. 1977). in the region of lowest mean salinity inhabited (sorgenfrei 1958, table 11), for 5 species from the baltic during the eemian, i.e. modiolus modiolus, turritella communis, nucula nitidosa, nucula nucleus, and spisula subtruncata, a salinity between 20‰ and 25‰ can be shown, which is above the present conditions in the baltic. however, this might be caused by a higher tide in inner danish waters during the eemian, in this way giving the same situation as known during the atlantic, when the salinity was higher. unlike the situation in the bælt sea, the present-day characterising species macoma balthica is recorded from the localities on sjælland, strandegaards dyrehave and møn. the drop in number of species among bivalves and gastropods from the bælt sea to the baltic is from 45 to 11 species respectively in the present-day fauna (sorgenfrei 1958), which is of the same order of magnitude as seen in the fossil fauna from the bælt sea to the baltic during the eemian. this means that in some waywe have to do with the same basin structure/hydrographic situation, although the salinity was higher during the eemian in the innermost danish waters than at present. the kattegat age: eemian climatic regions: asbl class bivalvia subclass heterodonta order myoida hiatella arctica (linnaeus 1758) total for climatic regions asbl: 1 (3.1%) climatic regions: .sbl class gastropoda subclass prosobranchia order archaeogastropoda scissurella crispata fleming 1828 order neogastropoda buccinum undatum linnaeus 1758 class bivalvia subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 total for climatic regions .sbl: 3 (9.4%) climatic regions: ..bl class gastropoda subclass prosobranchia order neotaenioglossa littorina littorea (linnaeus 1758) hydrobia ulvae (pennant 1777) rissoa inconspicua alder 1844 rissoa violacea desmarest 1814 bittium reticulatum (da costa 1778) turritella communis risso 1826 lunatia alderi (forbes 1838) order neogastropoda hinia pygmaea (lamarck 1822) hinia reticulata (linnaeus 1758) subclass heterobranchia order heterostropha chrysallida spiralis (montagu 1803) turbonilla lactea (linné 1758) class bivalvia subclass pteriomorpha order pterioida ostrea edulis linnaeus 1758 subclass heterodonta order veneroida acanthocardia echinata (linnaeus 1758) parvicardium exiguum (gmelin 1791) cerastoderma edule (linnaeus 1758) geus bulletin no 3.pmd 28-06-2004, 08:45118 119 laevicardium crassum (gmelin 1791) spisula solida (linnaeus 1758) spisula subtruncata (da costa 1778) scrobicularia plana (da costa 1778) abra prismatica (montagu 1803) paphia aurea (gmelin 1791) timoclea ovata (pennant 1777) order myoida corbula gibba (olivi 1792) total for climatic regions ..bl: 23 (71.9%) climatic regions: ...l class gastropoda subclass opisthobranchia order bullomorpha haminoea navicula (da costa 1778) class bivalvia subclass pteriomorpha order mytiloida mytilaster lineatus (gmelin 1791) subclass heterodonta order veneroida lucinella divaricata (linnaeus 1758) gastrana fragilis (linnaeus 1758) paphia senescens (cocconi 1873) total for climatic regions ...l: 5 (15.6%) the eemian kattegat: 32 (13.0%) from the region of the kattegat sensu lato area, including the bordering landmasses with fjords, sounds and the northern part of the lillebælt, storebælt, and øresund, three kinds of localities have been met with. as the information obtained from this area is based on a very different kind of material, the localities will be treated apart. the first locality to be considered is ejby bro in northern sjælland on the isefjord. here erik rasmussen (in madsen 1968) described an in situ marine deposit. out of 15 mollusc species, two are lusitanian, viz. lucinella divaricata and paphia aurea senescens, both of which are characteristic fossils of the eemian. furthermore, out of ten boreo-lusitanian species, seven are not recorded from the recent isefjord, viz. rissoa violacea, lunatia alderi, hinia reticulata (the british form (fretter & graham 1984, p. 495)), ostrea edulis, laevicardium crassum, paphia aurea, and spisula solida. among these are also the less tolerant species regarding the region of lowest mean salinity inhabited, which is five species between 28‰ and 33‰. the present-day figure of salinity for the isefjord is between 18 and 20‰. although we have to do with a saalien glacial topography, we may explain the higher salinity as a result of a higher tidal amplitude during the eemian. only two species occurring in the ejby bro locality extend into the subarctic: buccinum undatum and mytilus edulis, besides one – hiatella arctica – having a wide range. all are present also in the recent isefjord waters. this fauna points to a more oceanic environment than today and with a higher temperature; furthermore, the finds point to shallow-water or even beach deposits with some tidal influence (rasmussen in madsen 1968). the second kind of locality includes redeposited sediments: either floes in the glacial deposits like the stautrup locality at aarhus or fluvioglacial deposits as at høng in western sjælland (nordmann 1928, pp. 64– 65; ødum 1933; sorgenfrei 1945). both of these localities carried some of the characteristic molluscs of the eemian deposits. from høng the following have been recorded: lucinella divaricata and paphia aurea senescens together with 11 species known from the boreo-lusitanian zone and mytilus edulis known from the subarctic to the lusitanian regions. only four species were also found at the ejby bro locality, and the ‘new’ ones (nine species) point to deeper water with such species as acanthocardia echinata, corbula gibba and turritella communis. the stautrup material records new species to the kattegat region such as rissoa inconspicua, mytilaster lineatus, gastrana fragilis, abra prismatica, chrysallida spiralis, and haminoea navicula, all except mytilaster lineatus and gastrana fragilis being boreo-lusitanian, while the two bivalves are limited to the lusitanian and characteristic of the eemian – together with paphia aurea senescens, which is also present in the stautrup floe, as discussed by sorgenfrei (1945). from the material hitherto discussed it appears that the eemian sea deposits were known from the kattegat region both in a shallow-water facies (ejby bro) and a deeper-water facies (høng), the latter only from redeposited material. therefore, it is of great importance from a palaeogeographical point of view that information now has been obtained from borings in the central part of the kattegat, on the island of anholt (lykke-andersen et al. 1993). foraminifera from this well (seidenkrantz 1993) revealed a marine upper saalian and eemian sequence. the macrofossils from the well were kindly placed at my disposal. it appears that a temperate fauna with species such as turritella geus bulletin no 3.pmd 28-06-2004, 08:45119 120 communis, aclis minor, hiatella arctica, nuculana minuta, and scissurella crispata is resting on an arctic deposit with portlandia arctica. the information that we do have is that a turritella facies within the kattegat region in the eemian found in an in situ position sustains the view of a continuation towards the north not only of the shallow-water deposits, but also of the deeper-water environment during the eemian. considering the bottom communities of the present-day kattegat as revealed by petersen (1913), turritella communis is found in different associations at depths from 12–19 to 35 m on sand, fine sand and clay. the species from the eemian of the kattegat associated with turritella communis in the kattegat of today are hiatella arctica, nuculana minuta, acanthocardia echinata, corbula gibba, and hinia pygmaea. considering the eemian faunas demonstrated from the kattegat region, it appears that here the lowest mean salinity inhabited by the species in question is within the present-day salinities reached for this area, at about 33‰, although the salinities for the present fjords bordering the kattegat have a lower salinity, as mentioned in the case of the isefjord with the ejby bro locality. the north sea age: eemian climatic regions: asb. class bivalvia subclass heterodonta order veneroida tridonta elliptica (brown 1827) total for climatic regions asb. : 1 (1.1%) climatic regions: asbl class gastropoda subclass prosobranchia order neotaenioglossa littorina saxatilis (olivi 1792) class bivalvia subclass heterodonta order veneroida thyasira flexuosa (montagu 1803) tridonta montagui (dillwyn 1817) order myoida mya truncata linnaeus 1758 hiatella arctica (linnaeus 1758) total for climatic regions asbl: 5 (5.5%) climatic regions: .sb. class bivalvia subclass heterodonta order veneroida arctica islandica (linnaeus 1767) total for climatic regions .sb. : 1 (1.1%) climatic regions: .sbl class gastropoda subclass prosobranchia order neotaenioglossa littorina obtusata (linnaeus 1758) lacuna vincta (montagu 1803) order neogastropoda buccinum undatum linnaeus 1758 class bivalvia subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 modiolus modiolus (linnaeus 1758) order pterioida heteranomia squamula (linnaeus 1758) subclass heterodonta order veneroida mysella bidentata (montagu 1803) tellimya ferruginosa (montagu 1803) parvicardium ovale (sowerby 1840) macoma balthica (linnaeus 1758) total for climatic regions .sbl: 10 (11.0%) climatic regions: ..bl class gastropoda subclass prosobranchia order archaeogastropoda gibbula cineraria (linnaeus 1758) order neotaenioglossa littorina littorea (linnaeus 1758) lacuna parva (montagu 1803) hydrobia ulvae (pennant 1777) onoba vitrea (montagu 1803) rissoa albella lovén 1846 rissoa inconspicua alder 1844 rissoa membranacea (j. adams 1800) rissoa parva (da costa 1779) rissoa violacea desmarest 1814 caecum glabrum (montagu 1803) bittium reticulatum (da costa 1778) turritella communis risso 1826 aporrhais pespelicani (linnaeus 1758) lunatia alderi (forbes 1838) order heterogastropoda geus bulletin no 3.pmd 28-06-2004, 08:45120 121 triphora adversa (montagu 1803) cerithiopsis tubercularis (montagu 1803) epitonium clathrus (linnaeus 1758) order neogastropoda hinia pygmaea (lamarck 1822) hinia reticulata (linnaeus 1758) cytharella coarctata (forbes 1840) subclass heterobranchia order heterostropha brachystomia eulimoides hanley 1844 odostomia scalaris macgillivray 1843 chrysallida obtusa (brown 1827) chrysallida spiralis (montagu 1803) ebala nitidissima (montagu 1803) odostomia albella lovén 1846 turbonilla crenata (brown 1827) turbonilla lactea (linné 1758) subclass opisthobranchia order bullomorpha acteon tornatilis (linnaeus 1758) philine aperta (linnaeus 1767) order anaspidea retusa truncatula (bruguière 1792) akera bullata müller 1776 class bivalvia subclass palaeotaxodonta order nuculoida nucula nitidosa winckworth 1930 nucula sulcata (bronn 1831) subclass pteriomorpha order mytiloida modiolula phaseolina (philippi 1844) modiolaria tumida (hanley 1843) order pterioida aequipecten opercularis (linnaeus 1758) chlamys varia (linnaeus 1758) ostrea edulis linnaeus 1758 subclass heterodonta order veneroida lepton nitidum (turton 1822) acanthocardia echinata (linnaeus 1758) parvicardium exiguum (gmelin 1791) parvicardium scabrum (philippi 1844) cerastoderma edule (linnaeus 1758) spisula subtruncata (da costa 1778) ensis ensis (linnaeus 1758) phaxas pellucidus (pennant 1777) angulus tenuis (da costa 1778) tellina donacina linnaeus 1758 fabulina fabula (gmelin 1791) donax vittatus (da costa 1778) scrobicularia plana (da costa 1778) abra alba (wood 1802) abra prismatica (montagu 1803) chamelea striatula (da costa 1778) tapes decussatus (linnaeus 1758) timoclea ovata (pennant 1777) venerupis pullastra (montagu 1803) dosinia lincta (montagu 1803) mysia undata (pennant 1777) order myoida corbula gibba (olivi 1792) saxicavella jeffreysi winckworth 1930 barnea candida (linnaeus 1758) subclass anomalodesmata order pholadomyoida thracia phaseolina (lamarck 1818) thracia villosiuscula (macgillivray 1827) total for climatic regions ..bl: 66 (72.5%) climatic regions: ...l class gastropoda subclass opisthobranchia order bullomorpha haminoea navicula (da costa 1778) class bivalvia subclass pteriomorpha order mytiloida mytilaster lineatus (gmelin 1791) subclass heterodonta order veneroida lucinella divaricata (linnaeus 1758) plagiocardium papillosum poli 1795 gastrana fragilis (linnaeus 1758) abra segmentum (récluz 1843) paphia senescens (cocconi 1873) gouldia minima (montagu 1803) total for climatic regions ...l: 8 (8.8%) the eemian north sea: 91 (36.8%) regarding the eemian deposits of south-western jylland several localities are included: tønder and surroundings, forballum, farup, ydre bjergum, mandø hølade (many places) and inder bjergum. furthermore, also molluscs of eemian age are recorded from the map sheet blaavands huk (fig. 1). in all, 91 molluscan species have been recorded, 53 bivalves and 38 gastropods. this high number of species comprises 32 species new to the eemian compared to the bælt sea deposits. geus bulletin no 3.pmd 28-06-2004, 08:45121 122 considering the new elements from the climatic point of view, two are lusitanian species, plagiocardium papillosum and mytilaster lineatus, which are regarded as part of the characteristic species of the eemian fauna (nordmann 1928). by far the largest group of new species in the north sea deposits are the boreo-lusitanian. only one – tridonta elliptica – does not extend into the lusitanian region. this is close to the situation found in the bælt sea region, where only two species; arctica islandica and zirfaea crispata, do not reach the lusitanian region. arctica islandica – so common in the bælt sea deposits – has been recorded only in a single find of a juvenile specimen at mandø hølade in south-western jylland. the other species, zirfaea crispata, has not been demonstrated at all in the danish north sea eemian deposits. the boreo-lusitanian species mactra stultorum has a high salinity requirement, occurring in the bælt sea region but not in the danish north sea region during the eemian. this has been used as an argument against uniting in time the deposits found on the western and eastern sides of southern jylland. “on a voulu y voir une preuve que ces deux bassins de mer n’out en réalité rien eu à faire l’un avec l’autre, en sorte qu’ils pourraient très bien être d’àges fort différents” (nordmann 1928, p. 63). when the 15 mollusc species of the eemian bælt sea fauna mentioned earlier with a salinity requirement between 30–33‰ are remembered, the former focus on mactra stultorum is of less significance. while nordmann states that the fauna on both sides of the jylland peninsula can be regarded as one, he is right from the point of view of climatic conditions, as demonstrated above. however, he is also arguing for a connection between the north sea and the bælt sea (nordmann 1928, p. 63): “les passes entre les parties orientale et occidentale de la mer eemienne, ce qui, autrement parlant, signifie les vallées et les plaines entre les collines insulaires qui sont aujourd’hui occupées par les plaines de landes du slesvig et du holstein, sont sans doute très etroites”. however, it is the author’s opinion that the difference between the southern danish localities to the east (the bælt sea) and to the west (the north sea) can be regarded as differences in facies that are also found in present-day danish waters. here the occurrences of donax vittatus in the eemian north sea deposits, but not in the bælt sea deposits, should be considered. donax vittatus is a characteristic species of the high-energy coastal environment, where it is found all along the present-day west coast, but not in the inner danish waters from the northernmost part of the east coast of jylland. however, many of the other molluscs might have made their way to the bælt sea and the baltic region if there has been only the slightest passage over southern jylland. in the case of a passage, the situation can be looked upon as a parallel to the present-day marine colonisation of the limfjord after the breaking through by the north sea at the spit at agger in 1825. at the end of the 19th century a rich marine fauna could be recorded in the limfjord (collin 1884; petersen 1888). the question of connection between the north sea and the bælt sea over the southern part of the peninsula of jylland can be considered, based on the mollusc assemblages. it appears that the difference in faunal composition during the eemian in both seas was very similar to the difference in faunas during the holocene, although with lusitanian shallow-water species occurring during the eemian. therefore it must be concluded that the land–sea configuration must be very much the same during the two periods so jylland was also a peninsula during the eemian. vendsyssel age: eemian climatic regions: asb. class gastropoda subclass prosobranchia order neogastropoda oenopota incisula (verrill 1882) oenopota violacea (mighels & adams 1842) bela exarata g.o. sars 1818 class bivalvia subclass palaeotaxodonta order nuculoida nuculana pernula (müller 1776) yoldiella lenticula (möller 1842) subclass pteriomorpha order mytiloida musculus niger (gray 1824) subclass heterodonta order veneroida clinocardium ciliatum (fabricius 1780) serripes groenlandicus (bruguière 1798) macoma calcarea (gmelin 1791) total for climatic regions asb. : 9 (16.4%) geus bulletin no 3.pmd 28-06-2004, 08:45122 123 climatic regions: asbl class gastropoda subclass prosobranchia order neogastropoda oenopota trevelliana (turton 1834) subclass opisthobranchia order thecosomata limacina retroversa (fleming 1823) class bivalvia subclass palaeotaxodonta order nuculoida nuculoma tenuis (montagu 1808) yoldiella frigida (torell 1859) subclass heterodonta order veneroida leptaxinus ferruginosus (forbes 1844) order myoida mya truncata linnaeus 1758 hiatella arctica (linnaeus 1758) total for climatic regions asbl: 7 (12.7%) climatic regions: .sb. class bivalvia subclass palaeotaxodonta order nuculoida nuculana minuta (müller 1776) total for climatic regions .sb. : 1 (1.8%) climatic regions: .sbl class gastropoda subclass prosobranchia order neotaenioglossa lacuna vincta (montagu 1803) order neogastropoda buccinum undatum linnaeus 1758 class bivalvia subclass palaeotaxodonta order nuculoida yoldiella lucida (lovén 1846) subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 subclass heterodonta order veneroida mysella bidentata (montagu 1803) parvicardium ovale (sowerby 1840) total for climatic regions .sbl: 6 (10.9%) climatic regions: ..b. class gastropoda subclass heterobranchia order heterostropha chrysallida eximia (jeffreys 1849) total for climatic regions ..b. : 1 (1.8%) climatic regions: ..bl class gastropoda subclass prosobranchia order neotaenioglossa alvania abyssicola (forbes 1850) rissoa parva (da costa 1779) bittium reticulatum (da costa 1778) turritella communis risso 1826 aporrhais pespelicani (linnaeus 1758) lunatia alderi (forbes 1838) order heterogastropoda vitreolina philippii (rayneval & ponzi1854) order neogastropoda hinia incrassata (ström 1768) hinia reticulata (linnaeus 1758) mangelia brachystoma (philippi 1844) raphitoma linearis (montagu 1803) subclass heterobranchia order heterostropha eulimella scillae (scacchi 1835) ondina divisa (j. adams 1797) odostomia turrita hanley 1844 subclass opisthobranchia order bullomorpha acteon tornatilis (linnaeus 1758) order anaspidea retusa umbilicata (montagu 1803) class bivalvia subclass palaeotaxodonta order nuculoida nucula nucleus (linnaeus 1767) nucula sulcata (bronn 1831) yoldiella philippiana (nyst 1845) subclass pteriomorpha order pterioida pseudamussiumseptemradiatum (müller 1776) similipecten similis (laskey 1811) subclass heterodonta order veneroida acanthocardia echinata (linnaeus 1758) parvicardium minimum (philippi 1836) ensis ensis (linnaeus 1758) phaxas pellucidus (pennant 1777) abra alba (wood 1802) abra nitida (müller 1776) abra prismatica (montagu 1803) geus bulletin no 3.pmd 28-06-2004, 08:46123 124 kelliella miliaris (philippi 1844) chamelea striatula (da costa 1778) order myoida corbula gibba (olivi 1792) total for climatic regions ..bl: 31 (56.4%) the eemian vendsyssel: 55 (22.3%) in vendsyssel the mollusc faunas in the skærumhede sequence have been studied by nordmann (jessen et al. 1910) and petersen (bahnson et al. 1974). the conclusion reached in the latter study on molluscs points out that the difference between the boreo-lusitanian community – the turritella terebra zone – in the boring and the typical eemian community in southern denmark is a difference in facies. this statement will be discussed now on the basis of all the mollusc found and listed according to their climatic regions. among the 55 species of molluscs recorded from the eemian in the vendsyssel region, no lusitanian species occur, but boreo-lusitanian species count for more than half of the assemblages (31 species, 56.4%). of these, 25 are mentioned in the list for region of lowest mean salinity inhabited built on information from the danish waters in the transition area between the north sea (at esbjerg) and the baltic (gulf of bothnia) (sorgenfrei 1958). nearly 4/5 of this number have their region of lowest mean salinity between 24‰ and 34‰, which is the minimum, and mean salinities at the passage belt between the north sea – skagerrak and e and nw kattegat. this situation for an eemian assemblage indicates a high degree of similarity with the present-day environment in this area. within the boreo-lusitanian group of molluscs there is a clear dominance of species belonging to the deeperwater environment at around the 100 m depth. however, there are some few species which belong at a depth of less than 20–30 m. such species are rissoa parva, hinia reticulata and retusa umbilicata, which within their depth range live in great abundance, which is not the case in the skærumhede sequences. therefore they can be regarded as allochthonous or as stray finds outside their environment. nordmann (in jessen et al. 1910) mentioned that rissoa parva and hinia reticulata were redeposited. he also mentioned bittium reticulatum, which according to the literature has some records from the deeper water (out to 250 m deep). taking the whole group of boreo-lusitanian species, there are many which occur in the tidal and shallow-water environments but have a wide range of depth. rather few are those which are mainly connected with the deeper water – here depths of more than 10 m – viz.: turritella communis, lunatia alderi, mangelia brachystoma, raphitoma linearis, eulimella scillae, nucula sulcata, pseudamussium septemradiatum and kelliella miliaris. the last one strengthens the depth indications to be more than 100 m, since the species is recently known only from the deeper part of the skagerrak. chrysallida eximia, which is found at depths from 20–1000 m is here regarded as a boreal species, although it extends into the northernmost part of the lusitanian zone, but there in deeper water. among the species known from the subarctic to the lusitanian regions, the finds of mytilus edulis (jessen et al. 1910; bahnson et al. 1974) can be regarded from the same point of view as mentioned in connection with rissoa parva and hinia reticulata. the few finds of mytilus specimens clearly indicate that this normally gregariously living species is found outside its living zone and can be regarded also as stray finds. the other finds of species occurring in the subarctic to the lusitanian zone do not oppose the view of a deeper-water environment. taking the species together which do not occur in the lusitanian zone but extend into the arctic or subarctic (nuculana minuta), there are two species clinocardium ciliatum and serripes groenlandicus, which do pose a problem regarding their climatic indications within the skærumhede sequence. these two species are found only in the high boreal zone (norway north of lofoten). nordmann (jessen et al. 1910, pp. 127– 128) mentioned the occurrences of these species off iceland together with other species with which they occur in the skærumhede sequence. this picture has been to some degree changed by the publication of parts of the zoology of iceland, so it appears that turritella communis is no longer regarded as being part of the icelandic fauna, while eulimella scillae (only empty shells) might be added. lunatia alderi and raphitoma linearis mentioned by nordmann are confirmed (thorson 1941), and so is acteon tornatilis (lemche 1938), in the new literature. however, in this way we still face the question of the climatic conditions indicated by the molluscan assemblage. however, as seen from the latest skærumhede boring (bahnson et al. 1974, fig. 7), there is a clear transition zone between the turritella communis zone and the establishment of the turritella erosa zone. in this zone, the so-called abra nitida zone, we still find the geus bulletin no 3.pmd 28-06-2004, 08:46124 125 mixing of species with different climatic affinities, just as we do not have a sharp border zone to tell where we actually leave the eemian and pass into the weichselian. but we do have a well-defined bottom community for what we must call the eemian from the vendsyssel region, and that is the turritella communis community. this community we find in deeper water, also in the recent danish waters. oenopota incisula is one of the species also occurring in the transition zone between the two turritella communities, the boreal–lusitanian with t. communis, and the arctic with t. erosa. oenopota incisula has been found in both borings (jessen et al. 1910; bahnson et al. 1974, fig. 7), but the species has still not been recorded from the european coast of the north atlantic in recent time. in this way oenopota incisula becomes one of the few species extinct in our part of the world since the early/middle weichselian. finally the six species with a wide range within the climatic regions are all found in deeper water. however, here yoldiella frigida occurs in deeper water in the southern part of its range. limacina retroversa is a pelagic species and has been recorded in the present day to penetrate into the kattegat – bælt sea regions. skagen age: eemian climatic regions: asb. class bivalvia subclass palaeotaxodonta order nuculoida nuculana pernula (müller 1776) total for climatic regions asb. : 1 (7.1%) climatic regions: asbl class gastropoda subclass opisthobranchia order thecosomata limacina retroversa (fleming 1823) class scaphopoda siphonodentaliumlobatum (sowerby 1860) class bivalvia subclass palaeotaxodonta order nuculoida yoldiella frigida (torell 1859) subclass heterodonta order myoida hiatella arctica (linnaeus 1758) total for climatic regions asbl: 4 (28.6%) climatic regions: .sbl class scaphopoda antalis entalis (linnaeus 1758) class bivalvia subclass pteriomorpha order pterioida delectopecten vitreus (gmelin 1791) total for climatic regions .sbl: 2 (14.3%) climatic regions: ..bl class gastropoda subclass prosobranchia order neotaenioglossa lunatia alderi (forbes 1838) subclass opisthobranchia order bullomorpha philine catena (montagu 1803) class scaphopoda cadulus subfusiforme (m. sars 1865) entalina tetragona (brocchi 1814) class bivalvia subclass heterodonta order veneroida kelliella miliaris (philippi 1844) subclass anomalodesmata order pholadomyoida cochlodesma praetenue (pulteney 1799) total for climatic regions ..bl: 6 (42.9%) climatic regions: ...l class scaphopoda dentalium vulgare da costa 1778 total for climatic regions ...l: 1 (7.1%) the eemian skagen: 14 (5.7%) in the eemian part of the skagen well, one species, dentalium vulgare, mainly occurs within the lusitanian region and is not recorded from recent danish waters, although it is found in the southern part of the north sea. among the six boreo-lusitanian mollusc species, cochlodesma praetenue is rare in danish waters, while all the others as far as the information on habitats goes are connected with the deeper-water environment. this is also true for the species including the subarctic region, viz.: antalis entalis and delectopecten vitreus, where the latter in general has a depth range from 30– 600 m, but in skagerrak is found between 400 and 600 m. also antalis entalis has a wide range of depth, geus bulletin no 3.pmd 28-06-2004, 08:46125 126 but within danish water it is recorded only from 20–400 m. the four species: siphonodentalium lobatum, limacina retroversa, yoldiella frigida and hiatella arctica with a wide geographical distribution from the arctic to the lusitanian also have a wide range of depth. also nuculana pernula, here listed from the arctic to the boreal, could be considered to have a wide range like the other species mentioned above, however, in sw europe it is found only at depths greater than 400 m. in recent danish waters it occurs at depths from 20 to 200 m. seen together with the other species in this region during the eemian with a more southern affinity, nuculana pernula shows accordance, considering that the species in the arctic is mostly littoral. comparing the assemblages from the faunal elements from skærumhede, it appears that the skagen well depicts a deeper-water community without any influence from more shallow-water facies. the occurrence of three of the four scaphopoda cadulus subfusiforme, entalina tetragona and antalis entalis speaks in favour of an environment which is likely to be found in the deeper skagerrak, such as the amphilepis norvegica/pecten vitreus community with entalina tetragonaand of which one of the other molluscs, delectopecten vitreus, is considered a characteristic species. the demonstrated eemian assemblages in most of the danish regions show differences in climatic affinities – more lusitanian in the southern part – which, however, can be explained through an analysis of the faunas in their relation to depth – shallow-water lusitanian species in the south – and in some parts in relation to the community. the community concept which was developed for the recent danish waters does not find its equivalent in the eemian bælt sea and baltic regions, only partly in the kattegat region, but has a good correlation with the occurrences of the faunal assemblages in the regions to the north in jylland – the vendsyssel and skagen regions – only in these two regions does the succession of strata allow us to follow the development into the weichselian cooler/arctic molluscan fauna, although we find some deposits from the kattegat region (holmstrup and holbæk sites) which can be correlated to the early/middle weichselian in the vendsyssel and skagen region. early/middle weichselian species sorted after climatic affinities the kattegat age: early/middle weichselian climatic regions: a... class bivalvia subclass palaeotaxodonta order nuculoida portlandia arctica (gray 1824) total for climatic regions a... : 1 (14.3%) climatic regions: as.. class gastropoda subclass opisthobranchia order bullomorpha cylichna occulta (mighels 1841) total for climatic regions as.. : 1 (14.3%) climatic regions: asb. class bivalvia subclass palaeotaxodonta order nuculoida nuculana pernula (müller 1776) subclass heterodonta order veneroida macoma calcarea (gmelin 1791) total for climatic regions asb. : 2 (28.6%) climatic regions: asbl class bivalvia subclass palaeotaxodonta order nuculoida nuculoma tenuis (montagu 1808) subclass heterodonta order myoida mya truncata linnaeus 1758 hiatella arctica (linnaeus 1758) total for climatic regions asbl: 3 (42.9%) the early/middle weichselian kattegat: 7 (2.8%) in the kattegat region, two localities (holbæk and holmstrup, fig. 1) have mollusc faunas of early/middle weichselian age, which could be correlated to the older yoldia clay deposits found in the vendsyssel and skagen regions. nordmann (ødum 1933) described the mollusc fauna in the borings at holbæk to represent part of the portgeus bulletin no 3.pmd 28-06-2004, 08:46126 127 landia arctica zone in the skærumhede sequence (jessen et al. 1910). later, petersen & buch (1974) referred the outcrops at holmstrup with marine clay, characterised by macoma calcarea, to the weichselian part of the sequence close to the macoma calcarea zone (bahnson et al. 1974) in the new well at skærumhede in the vendsyssel region. the later correlation was mainly based on the foraminiferal studies (buch in petersen & buch 1974). also aminostratigraphic investigations have to some extent sustained this correlation (miller & mangerud 1985, p. 261). the mollusc faunas from these two localities point to an arctic environment as seen from the climatic indications, in that all seven species are found in the arctic and only three of them with a wide range: hiatella arctica, mya truncata and nuculoma tenuis. taken into account that these deposits show an arctic affinity, the interpretation of depth ranges of the species found point to the more shallow-water environment. this is also true for species such as nuculoma tenuis, which is recorded from offshore down to 300 m deep, but in the arctic is more littoral. macoma calcarea is found intertidal to several hundred metres, but only in deeper water in the southern part of the range for this species. in the arctic it is the characterising mollusc in shallow water: the arctic macoma calcarea community. furthermore, nuculana pernula, as mentioned earlier, is littoral in the arctic. the conclusion to draw from these finds in the kattegat region is that we have a part – the more shallow water – of the arctic zones recorded from the vendsyssel region represented within the kattegat region – southern part. vendsyssel age: early/middle weichselian climatic regions: a... class bivalvia subclass palaeotaxodonta order nuculoida portlandia arctica (gray 1824) total for climatic regions a... : 1 (2.8%) climatic regions: as.. class gastropoda subclass prosobranchia order neotaenioglossa alvania cruenta odhner 1915 turritella erosa couthouy 1838 subclass opisthobranchia order bullomorpha cylichna occulta (mighels 1841) class bivalvia subclass pteriomorpha order arcoida bathyarca glacialis (gray 1824) subclass anomalodesmata order pholadomyoida pandora glacialis leach 1819 lyonsia arenosa (möller 1842) total for climatic regions as.. : 6 (16.7%) climatic regions: asb. class gastropoda subclass prosobranchia order neotaenioglossa alvania scrobiculata (möller 1842) alvania jan mayeni (friele 1886) lunatia pallida (broderip & sowerby 1829) order neogastropoda oenopota incisula (verrill 1882) admete viridula (fabricius 1780) subclass opisthobranchia order anaspidea retusa obtusa (montagu 1803) class bivalvia subclass palaeotaxodonta order nuculoida nuculana pernula (müller 1776) yoldia hyperborea lovén 1859 yoldiella lenticula (möller 1842) subclass pteriomorpha order mytiloida musculus laevigatus (gray 1824) musculus niger (gray 1824) crenella decussata (montagu 1803) order pterioida palliolum greenlandicum (sowerby 1842) subclass heterodonta order veneroida axinopsida orbiculata (g.o. sars 1878) tridonta borealis schumacher 1817 tridonta elliptica (brown 1827) clinocardium ciliatum (fabricius 1780) serripes groenlandicus (bruguière 1798) macoma calcarea (gmelin 1791) total for climatic regions asb. : 19 (52.8%) geus bulletin no 3.pmd 28-06-2004, 08:46127 128 climatic regions: asbl class gastropoda subclass prosobranchia order neotaenioglossa natica affinis (gmelin 1790) subclass opisthobranchia order bullomorpha cylichna alba (brown 1827) class bivalvia subclass palaeotaxodonta order nuculoida nuculoma tenuis (montagu 1808) yoldiella frigida (torell 1859) subclass heterodonta order veneroida tridonta montagui (dillwyn 1817) order myoida mya truncata linnaeus 1758 hiatella arctica (linnaeus 1758) total for climatic regions asbl: 7 (19.4%) climatic regions: .sb. class bivalvia subclass pteriomorpha order pterioida chlamys islandica (o.f. müller 1776) subclass heterodonta order myoida panomya arctica (lamarck 1818) total for climatic regions .sb. : 2 (5.6%) climatic regions: .sbl class gastropoda subclass prosobranchia order neotaenioglossa lacuna vincta (montagu 1803) total for climatic regions .sbl: 1 (2.8%) the early/middle weichselian vendsyssel: 36 (14.6%) the mollusc fauna from the vendsyssel region during the early/middle weichselian amounts to 36 species, which would have been even more if not reduced to this number by excluding species regarded as redeposited by v. nordmann (jessen et al. 1910). almost all the recorded species have been found in the skærumhede borings i and ii (jessen et al. 1910; bahnson et al. 1974 respectively), except lunatia pallida, bathyarca glacialis, musculus laevigatus and tridonta borealis, which have been recorded from the older yoldia clay elsewhere in vendsyssel. taken together, the two borings form a most excellent base for evaluating the mollusc faunal development in the early/middle weichselian represented by the arctic sea deposits characterised by portlandia arctica. however, on the basis of the material from skærumhede (bahnson et al. 1974), the portlandia arctica zone is divided into three parts each characterised by other macrofossils from the older to the younger beds: the turritella erosa, balanus crenata and macoma calcarea zones. by doing so, it is emphasised that the development in the arctic part of the marine sequence goes from a deeper-water facies into a shallow-water facies, where in the latter the macoma calcarea species is the dominant bivalve, as it is in the present-day arctic macoma community of east greenland (thorson 1933). from a climatic point of view, nearly all the molluscan species can be found in the high arctic, except lacuna vincta, panomya arctica and chlamys islandica, which are recorded only from the subarctic. however, two of them occur in the turritella erosa zone, which at the same time has the most abundant representation of the other species with northern/arctic affinities, including the purely arctic species portlandia arctica. therefore, no rise in temperature can be suggested on the basis of the molluscan record, only changes in facies through time. together with the dominant molluscan species portlandia arctica and macoma calcarea in the upper part of the arctic sequence – the balanus crenata and the macoma calcarea zones (as seen on fig. 7 in bahnson et al. 1974) – the following mollusc species have been recorded only from these zones: natica affinis, musculus niger, palliolum greenlandicum, and axinopsida orbiculata. all of them can be found in shallow water in the arctic. considering the change of facies from deeper water to shallow water and the lack of climatic changes as seen in the molluscan fauna the question arises of what length of time this development covers. thisquestionhasbeenansweredby the twoamsdates of the topmost part of the marine skærumhede sequence 33 m b.s., which give the age of around 32 000 14c years before present (aar-1410: 32 400 ± 520 and aar-1411: 32 050 ± 420 – both reservoir corrected 14c age (b.p.)). this shows that the marine arctic deposits in this part of the danish area represent nearly the whole part of the early and middle weichselian, because there is an unbroken marine sequence below the level for these ams dates and back into the eemian. on the basis of the correlation of the holmstrup sequence of clay with macoma calcarea, there are reageus bulletin no 3.pmd 28-06-2004, 08:46128 129 sons to think that also the southern part of the kattegat region was part of the older yoldia clay sea far into the weichselian, even though two ams datings from the holmstrup site were infinite (aar-1408: > 38 000 14c age (b.p.) and aar-1409: > 42 000 14c age (b.p.)). skagen age: early/middle weichselian climatic regions: a... class bivalvia subclass palaeotaxodonta order nuculoida portlandia arctica (gray 1824) total for climatic regions a... : 1 (25.0%) climatic regions: asb. class bivalvia subclass palaeotaxodonta order nuculoida nuculana pernula (müller 1776) yoldia hyperborea lovén 1859 subclass pteriomorpha order pterioida palliolum greenlandicum (sowerby 1842) total for climatic regions asb. : 3 (75.0%) the early/middle weichselian skagen: 4 (1.6%) in the skagen region, where all the information coming from the well has been described in more detail earlier, the macrofossil fauna can be presented on the basis of quantitative analyses and sedimentological data (appendix 3). therefore, the material can be seen on the background of a certain bottom community. such a relation was already seen realised in the discussion of the eemian strata in the skagen region, which pointed out that these strata could be correlated with an environment of the deeper part of the present skagerrak. however, the change found in the arctic section of the skagen well on the basis of the mollusc record occurs rather abruptly, turning the scenario into an arctic environment with ice-rafted minerogene material. such a palaeoenvironment is far from the communities demonstrated in recent danish waters, but is well known from east greenland. regarding the finds from the arctic part of the skagen sequence following the temperated eemian strata, it appears that there are no finds of bias from nearshore or indications of other climatic conditions than from the arctic. the four molluscan species found are recorded either only in the high-arctic, viz. portlandia arctica, or in the three zones from the arctic to the boreal, viz. nuculana pernula, yoldiahyperborea, and palliolum greenlandicum. this is a different situation than found in the vendsyssel region, where the two skærumhede wells revealed a clear transition zone – the so-called abra nitida zone – and within the arctic part, with occurrence of molluscs which have been regarded as redeposited (nordmann in jessen et al. 1910), such as: mytilus edulis, pseudamussium septemradiatum, zirfaea crispata, and bittium reticulatum. it has been argued in the present paper in connection with the holocene strata from skagen that, due to the expired isostatic uplift since the weichselian glaciation, the actual depth below present sea level of the holocene beds can be regarded as representing the palaeodepth – taken into account the eustatic movements through time. if this is true, the older strata – the eemian and early weichselian – may also be in a position below present sea level, which could reflect their palaeodepth, here also taking into consideration the older eustatic situation and the question of consolidation. neotectonic movements might be the black horse together with the higher level of the sea during the eemian, max. 7–8 m above the present level. this shows that the transition zone – the abra nitida zone – in the skærumhede ii well is found at a present depth below sea level of between 78 and 88 m (bahnson et al. 1974, fig. 7), while the sharp boundary between the temperate and the arctic zone in the skagen well is at a depth of 180 m b.s. the sea level must have been the same for the two stations at this time, so the difference in depth must be around 100 m. therefore the difference in development of the mollusc fauna within the two sequences depends on depths. this might also explain the occurrences of redeposited molluscs from shallow water in the skærumhede sequence and not in the skagen sequence by a closer coastal environment. the further development of the marine arctic in the skagen region has been truncated by the glaciation within the area. however, as demonstrated earlier, the overlying glacigene deposits have indeed accumulated the ‘missing’ younger marine strata up to an age of around 32 000 before present, as shown by dating of the marine gases. geus bulletin no 3.pmd 28-06-2004, 08:46129 130 late weichselian species sorted after climatic affinities vendsyssel age: late weichselian climatic regions: a... class bivalvia subclass palaeotaxodonta order nuculoida portlandia arctica (gray 1824) subclass heterodonta order veneroida macoma torelli jensen 1904 total for climatic regions a... : 2 (5.7%) climatic regions: as.. class gastropoda subclass opisthobranchia order bullomorpha cylichna occulta (mighels 1841) class bivalvia subclass heterodonta order veneroida macoma loveni jensen 1904 subclass anomalodesmata order pholadomyoida pandora glacialis leach 1819 lyonsia arenosa (möller 1842) total for climatic regions as.. : 4 (11.4%) climatic regions: asb. class gastropoda subclass prosobranchia order neotaenioglossa lunatia pallida (broderip & sowerby 1829) order neogastropoda boreotrophon clathratus (linnaeus 1767) buccinum cyaneum bruguière 1792 neptunea despecta (linnaeus 1758) oenopota turricola (montagu 1803) subclass opisthobranchia order anaspidea retusa obtusa (montagu 1803) class bivalvia subclass palaeotaxodonta order nuculoida nuculana pernula (müller 1776) yoldiella lenticula (möller 1842) subclass pteriomorpha order mytiloida musculus laevigatus (gray 1824) musculus niger (gray 1824) subclass heterodonta order veneroida axinopsida orbiculata (g.o. sars 1878) tridonta borealis schumacher 1817 macoma calcarea (gmelin 1791) total for climatic regions asb. : 13 (37.1%) climatic regions: asbl class gastropoda subclass prosobranchia order neotaenioglossa littorina saxatilis (olivi 1792) natica affinis (gmelin 1790) subclass opisthobranchia order bullomorpha cylichna alba (brown 1827) order thecosomata limacina retroversa (fleming 1823) class bivalvia subclass palaeotaxodonta order nuculoida nuculoma tenuis (montagu 1808) subclass heterodonta order veneroida thyasira flexuosa (montagu 1803) order myoida mya truncata linnaeus 1758 hiatella arctica (linnaeus 1758) total for climatic regions asbl: 8 (22.9%) climatic regions: .sb. class bivalvia subclass pteriomorpha order pterioida chlamys islandica (o.f. müller 1776) subclass heterodonta order veneroida arctica islandica (linnaeus 1767) order myoida zirfaea crispata (linnaeus 1758) total for climatic regions .sb. : 3 (8.6%) climatic regions: .sbl class polyplacophora order neoloricata geus bulletin no 3.pmd 28-06-2004, 08:46130 131 tonicella marmorea (fabricius 1780) class gastropoda subclass prosobranchia order neotaenioglossa lacuna vincta (montagu 1803) order neogastropoda buccinum undatum linnaeus 1758 class bivalvia subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 subclass heterodonta order veneroida macoma balthica (linnaeus 1758) total for climatic regions .sbl: 5 (14.3%) the late weichselian vendsyssel: 35 (14.2%) marine molluscs from the time after the main glaciation of denmark have been recorded only from vendsyssel and skagen, but new studies are increasing our knowledge from the kattegat region, but are not included in this work. the marine mollusc assemblages in vendsyssel have been mainly based on open profiles. the recorded mollusc speciesamount to35.however, therearesomeproblems in classifying the whole fauna in communities. three climatic groups can be demonstrated: a purely arctic and subarctic with six species, an arctic, subarctic and boreal consisting of 13 species, and a group of species which does not enter the arctic but might be found extending into the lusitanian. a fourth group with a wide range can be differentiated according to depth ranges. littorina saxatilis is intertidal, while natica affinis, cylichna alba, and thyasira flexuosa are species found from infratidal to great depths. nuculoma tenuis is found offshore to 300 m, and limacina retroversa is pelagic. so the species with a climatically wide range show the existence of two facies: a littoral and a deeperwater facies. considering the five species with a representation within the subarctic, boreal and lusitanian regions, macoma balthica is a shallow-water species, mytilus edulis is eulittoral and here the species occurs in great quantities, lacuna vincta is intertidal to depths of 60 m, tonicella marmorea from 0 to 183 m but more common at depths of less than 20 m, and finally buccinum undatum which is found sublittorally to great depths (1200 m). the characteristic depth for the dominating part of these molluscs is seen to be the shallow water. two of the species from the subarctic–boreal group, arctica islandica and zirfaea crispata, have very different characteristics as to depth of living. the depth range for arctica islandica is in general intertidal to 480 m, but jensen (1902, pp. 38–39) writes that arctica is a genuine boreal species and bases this on the fact that in the white sea area, which is in the northernmost part of its distribution, it is found in more shallow water than elsewhere. a relatively high temperature is reached only in the shallow water in this region. this means that arctica islandica in the present setting among other molluscs of purely arctic and subarctic relations must be an indicator of boreal waters in the shallow-water environment. zirfaea crispata is on the other hand a clear indicator of shallow water, having its range of depth between the low tide line and out to a depth of about 7 m. the third member of the subarctic–boreal climatic region, chlamys islandica, is known from the tidal zone and down to depths of 300 m. of the 13 species represented within the climatic regions of the arctic, subarctic and boreal, half of the members are infralittoral from 6–10 m to great depths: lunatia pallida, boreotrophon clathratus, neptunea despecta, oenopota turricola, nuculana pernula, yoldiella lenticula, and musculus niger. the other half can be found in the tidal zone but also at greater depths. among these, macoma calcarea and tridonta borealis are the characteristic bivalves in the arctic shallowwater macoma calcarea community with the astarte borealis zone in the most shallow parts from 3 to about 12–14 m in east greenland (thorson 1933, pp. 8–18). according to thorson (1933), the astarte borealis zone is no tide-water community such as for instance the macoma balthica community in some boreal seas. this is discussed in further detail by madsen (1936, p. 71), who concludes: “the littoral fauna [north of c. 66–67°n lat. east greenland] is especially characterised by the absence of littoral molluscs and balanidea, notably mytilus edulis, littorina saxatilis var. groenlandica, and balanus balanoides, all of which occur south of the above-mentioned limit”. the arctic–subarctic species in vendsyssel, represented by four species, can be found at water depths from 2 to 5 m and out to around 200 m, except cylichna occulta, which has a depth range of from 20 m to nearly 400 m. the purely arctic species portlandia arctica and macoma torelli are recorded from 2 and 5 m out to geus bulletin no 3.pmd 28-06-2004, 08:46131 132 around 340 m and 90 m respectively. this shows that the arctic–subarctic part of the recorded species from the vendsyssel region have a wide range of occurrence restricted not by a single species to the shallowwater environment as is the case among the more temperate mollusca and the species with a wide climatic range: arctica islandica, zirfaea crispata, macoma balthica, mytilus edulis, and littorina saxatilis. it can therefore be stated that all the arctic species could be together in deeper water and some also in the more shallow water. however, some of the more temperate species are restricted to shallow water, and arctica islandica is a distinctly shallow-water species in the northern part of its range. so, within the time span of deposition for the late weichselian younger yoldia sea deposits, an amelioration of the shallowwater environment including the tidal zone must have happened. nordmann (1910) noticed that certain marine strata in the vendsyssel region had a distinct littoral fauna, and he suggested that these beds were of a more recent origin than the younger yoldia clay, introducing the so-called zirphaea transgression named after one of the characteristic bivalves from the shallow-water environment discussed above. however, petersen (1984), on the basis of molluscan studies combined with the many 14c dates, could conclude that the deposition of yoldia clay containing a cold marine fauna from deeper water is seen to continue into bølling, but with a contemporaneous temperate fauna in the shallow-water deposits – the zirphaea beds. the evaluation of the marine history from vendsyssel is highly influenced by experiences obtained from work on holocene raised marine deposits and recent bottom samples in east greenland, as seen in petersen (1986b, figs 2, 3), where beds from the astarte borealis zone with mytilus edulis are overlying the ophiocten zone (deeper part of the arctic macoma calcarea community (thorson 1933, pp. 18–27)) with portlandia arctica. these observations may also explain the occurrences of redeposited material mentioned within the discussion of the older yoldia clay at skærumhede in the vendsyssel region. still, it must be regarded as redeposited, but not necessarily differing so much in time, as long as the above-mentioned observations show that a more temperate zone can be found in the shallow-water environment contemporaneous with an arctic fauna in deeper water. skagen age: late weichselian climatic regions: a... class bivalvia subclass palaeotaxodonta order nuculoida portlandia arctica (gray 1824) total for climatic regions a... : 1 (11.1%) climatic regions: as.. class bivalvia subclass pteriomorpha order arcoida bathyarca glacialis (gray 1824) total for climatic regions as.. : 1 (11.1%) climatic regions: asb. class bivalvia subclass palaeotaxodonta order nuculoida nuculana pernula (müller 1776) yoldia hyperborea lovén 1859 yoldiella lenticula (möller 1842) total for climatic regions asb. : 3 (33.3%) climatic regions: asbl class gastropoda subclass opisthobranchia order gymnosomata clione limacina (phipps 1774) class scaphopoda siphonodentalium lobatum (sowerby 1860) class bivalvia subclass palaeotaxodonta order nuculoida yoldiella frigida (torell 1859) total for climatic regions asbl: 3 (33.3%) climatic regions: .sb. class bivalvia subclass palaeotaxodonta order nuculoida nuculana minuta (müller 1776) total for climatic regions .sb. : 1 (11.1%) the late weichselian skagen: 9 (3.6%) the skagen region has contributed with only nine species, out of which the clione limacina species must be taken with some reservation, being based on an geus bulletin no 3.pmd 28-06-2004, 08:46132 133 imprint only. all the recorded species have been taken from the skagen well core superjacent to the older eemian and weichselian deposits discussed above. therefore the fauna represents an assemblage from a certain depth – through time – and can be seen in relation to the sedimentological information (appendix 3). the granulometric composition in the older yoldia sea sequence reflects two maxima on the frequency curve,which tells that part of the material, other than the extremely fine-grained, can be taken as icerafted material. however, no such redeposited material was found in the fauna. the whole mollusc assemblage resembles the arca-astarte crenata community as described by thorson (1934) from hurry inlet, east greenland. first of all the arca glacialis (bathyarca glacialis) is represented among the molluscs recorded from the skagen well. this species is one of the characteristic species from this community. furthermore, the following species are mentioned (thorson 1934, p. 48): siphonodentalium vitreum (s. lobatum), leda pernula (nuculana p.), portlandia arctica, portlandia lenticula (yoldiella l.), portlandia frigida (yoldiella f.), and saxicava arctica (hiatella a.) which have all been recorded from the skagen well. only nuculana minuta, a subarctic–boreal species, and yoldia hyperborea are not recorded from hurry inlet. ockelmann (1958, pp. 19–22) mentioned that nuculana minuta is “lacking in the most high-arctic seas” and that according to thorson (1934) yoldia hyperborea is associated with calm, sheltered places, and besides yoldia hyperborea is otherwise known from only a few places in east greenland. these records from greenland demonstrate a good agreement with the danish late weichselian finds from the skagen well and permit further comparison with the arca-astarte crenata community. according to thorson (1933, p. 67), this community inhabits depths from ca. 45 to ca. 200 m. furthermore, it is poor in species, and the temperature is negative and constant all the year round. the well-dated strata of late weichselian age, from 15 000 – 10 000 before present in 14c years, around 17000 – 11 000 in calibrated age b.p. (see appendix 4) in the skagen well have no influence from the shallow-water boreo-arctic assemblage as found in the vendsyssel region around 13 000 before present (petersen 1984, p. 65), but reveal with their arctic deeperwater mollusc fauna a sharp boundary to the holocene boreo-lusitanian faunas. holocene species sorted after climatic affinities the boundary between the pleistocene and the holocene marine strata as seen in the skagen well is unique within the danish realm. furthermore, the approximately 115 m of holocene marine beds as described earlier represent a well-dated sequence to be compared to the other holocene marine finds recorded in the six regions, many of which offer dated strata as well. the following description will proceed in the same way as taken for the marine pleistocene strata, and it will present the entire molluscan fauna from each region from the point of view of climatic affinities for each species, with subsequent comments upon certain aspects for selected species. within each region, reference to well-dated strata and their molluscan assemblages will be given and worked out to facilitate the correlation in time to assemblages in other regions. finally, the juncture of occurrence during the holocene of some of the molluscan species in all regions can be estimated, as presented in appendix 6. the bælt sea climatic regions: asb. class gastropoda subclass opisthobranchia order anaspidea retusa obtusa (montagu 1803) class bivalvia subclass heterodonta order veneroida tridonta borealis schumacher 1817 total for climatic regions asb. : 2 (4.3%) climatic regions: asbl class gastropoda subclass prosobranchia order neotaenioglossa littorina saxatilis (olivi 1792) lacuna pallidula (da costa 1778) class bivalvia subclass pteriomorpha order mytiloida musculus discors (linnaeus 1767) subclass heterodonta order myoida mya truncata linnaeus 1758 hiatella arctica (linnaeus 1758) geus bulletin no 3.pmd 28-06-2004, 08:46133 134 total for climatic regions asbl: 5 (10.6%) climatic regions: .sb. class bivalvia subclass heterodonta order veneroida arctica islandica (linnaeus 1767) order myoida zirfaea crispata (linnaeus 1758) total for climatic regions .sb. : 2 (4.3%) climatic regions: .sbl class gastropoda subclass prosobranchia order neotaenioglossa littorina obtusata (linnaeus 1758) lacuna vincta (montagu 1803) onoba semicostata (montagu 1803) subclass heterobranchia order heterostropha omalogyra atomus (phillippi 1841) class bivalvia subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 modiolus modiolus (linnaeus 1758) subclass heterodonta order veneroida mysella bidentata (montagu 1803) parvicardium ovale (sowerby 1840) macoma balthica (linnaeus 1758) total for climatic regions .sbl: 9 (19.1%) climatic regions: ..b. class bivalvia subclass heterodonta order myoida mya arenaria linnaeus 1758 total for climatic regions ..b. : 1 (2.1%) climatic regions: ..bl class gastropoda subclass prosobranchia order neotaenioglossa littorina littorea (linnaeus 1758) littorina tenebrosa (montagu 1803) hydrobia ulvae (pennant 1777) hydrobia ventrosa (montagu 1803) rissoa albella lovén 1846 rissoa inconspicua alder 1844 rissoa membranacea (j. adams 1800) bittium reticulatum (da costa 1778) order heterogastropoda triphora adversa (montagu 1803) order neogastropoda hinia reticulata (linnaeus 1758) subclass heterobranchia order heterostropha odostomia conoidea winckworth 1932 subclass opisthobranchia order anaspidea retusa truncatula (bruguière 1792) akera bullata müller 1776 subclass pulmonata order basommatophora lymnaea peregra (müller 1774) class bivalvia subclass pteriomorpha order mytiloida modiolula phaseolina (philippi 1844) order pterioida subclass heterodonta order veneroida parvicardium exiguum (gmelin 1791) parvicardium scabrum (philippi 1844) cerastoderma edule (linnaeus 1758) spisula subtruncata (da costa 1778) angulus tenuis (da costa 1778) scrobicularia plana (da costa 1778) abra alba (wood 1802) paphia aurea (gmelin 1791) tapes decussatus (linnaeus 1758) venerupis pullastra (montagu 1803) order myoida corbula gibba (olivi 1792) barnea candida (linnaeus 1758) total for climatic regions ..bl: 28 (59.6%) the holocene bælt sea: 47 (19.0%) the holocene molluscs in the bælt sea region amount to 47 mainly boreo-lusitanian species. there are no purely lusitanian species, only two arctic–subarctic and boreal species, and five species with a wide range of distribution. this is a clear difference compared to the eemian molluscan fauna from the same region with no less than seven purely lusitanian species. this is not connected with differences in depth ranges. all of the subfossil species from the holocene could be found within the intertidal zone except: odostomia conoidea and parvicardium ovale occurgeus bulletin no 3.pmd 28-06-2004, 08:46134 135 ring below 5–10 m, or infratidal species such as tridonta borealis, modiolus modiolus and abra alba. abra alba characterises the bælt sea deeper-water community and tridonta borealis the community in the baltic, the so-called abra alba and astarte communities sensu petersen (1913, p. 16). the overall dominating part of the species could be associated with the macoma community sensu petersen (1913, p. 14) which nowadays is recorded from the fjords and the more sheltered coasts from the shore and out to a depth of 10–12 m. this community is named after macoma balthica, also present in the bælt sea subfossil molluscan fauna. the great similarity between the subfossil holocene fauna and the recent, as seen from the above-mentioned dates on the communities met with, is also revealed in the present distribution of the species in question. almost all the species occur in the present-day bælt sea or even further into the baltic region – east of darss. the few species no longer found in the bælt sea region are: littorina saxatilis, odostomia conoidea, modiolula phaseolina, ostrea edulis, parvicardium scabrum, angulus tenuis and venerupis pullastra which only extend into the kattegat region today, while paphia aurea, tapes decussatus, and omalogyra atomus are no longer recorded in the danish mollusc fauna. the absence of certain species is most likely an effect of the salinity reached in the bælt sea at present. however, as mentioned by rasmussen (1973, p. 303) in the case of venerupis pullastra, the distribution might be connected with the bottom conditions (see the chapter on molluscan species). most of the species now absent from the bælt sea were present during the atlantic, and, in all, 31 molluscan species have their first appearance in this period (appendix 6). among the 31 species with dated appearances in the atlantic, littorina saxatilis and mya truncata have a wide climatic range. arctica islandica is found in the boreal region, and lacuna vincta, onoba semicostata, mytilus edulis, mysella bidentata, and macoma balthica are subarctic–boreal–lusitanian. the rest of the species appearing in the atlantic amount to 23 species with dated appearances during the atlantic within the bælt sea area. the 31 species constitute 12.6% of the known subfossil molluscan finds from the late quaternary, while the total finds of molluscs from the holocene bælt sea represent 19.0% (47 species out of the 247 subfossil species). this might indicate that the more prolific fauna in the bælt sea was connected with a time interval when the inner danish waters were still affected by a higher tidal impact which expired at the beginning of the subboreal (petersen 1993). however, paphia aurea and ostrea edulis are still met with as food elements, although rather rare, in the iron age ‘køkkenmøddinger’ (kitchen middens) in the western part of the bælt sea (petersen 1985c, fig. 5, p. 22). mya arenaria has been found in the bælt sea embedded in strata from the atlantic. the appearance of this species within the atlantic strata is explained by the deep-burrowing habit of this infauna species. however, the first appearance of the species must be referred to the subatlantic after new dates in the bælt sea region following the study of the immigration of mya arenaria to danish waters (petersen et al. 1992b). the baltic climatic regions: .sbl class bivalvia subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 subclass heterodonta order veneroida macoma balthica (linnaeus 1758) total for climatic regions .sbl: 2 (10.5%) climatic regions: ..bl class gastropoda subclass prosobranchia order neotaenioglossa littorina littorea (linnaeus 1758) littorina tenebrosa (montagu 1803) hydrobia ulvae (pennant 1777) hydrobia ventrosa (montagu 1803) rissoa albella lovén 1846 rissoa inconspicua alder 1844 rissoa membranacea (j. adams 1800) bittium reticulatum (da costa 1778) aporrhais pespelicani (linnaeus 1758) order neogastropoda hinia reticulata (linnaeus 1758) subclass opisthobranchia order anaspidea retusa truncatula (bruguière 1792) subclass pulmonata order basommatophora lymnaea peregra (müller 1774) class bivalvia geus bulletin no 3.pmd 28-06-2004, 08:46135 136 subclass heterodonta order veneroida parvicardium exiguum (gmelin 1791) cerastoderma edule (linnaeus 1758) cerastoderma glaucum (poiret 1798) scrobicularia plana (da costa 1778) order myoida corbula gibba (olivi 1792) total for climatic regions ..bl: 17 (89.5%) the holocene baltic: 19 (7.7%) among the 19 species recorded from the holocene of the baltic, there is relatively many, which do not occur in the recent baltic fauna, viz.: rissoa albella, rissoa inconspicua, bittium reticulatum, hinia reticulata, retusa truncatula, and scrobicularia plana, but they are present in the neighbouring recent bælt sea fauna. furthermore, aporrhais pespelicani can be added which today extends into the kattegat region, while only empty shells have been recorded from the bælt sea region off kiel (arntz et al. 1976). so, although the spectacular oyster and tapes species did not occur as in the bælt sea, the subfossil baltic fauna has quite a few species that no longer live in the baltic. the geological mapping from this region (milthers 1908) does not provide further information on the chronostratigraphic position of these mollusc species, and is also rather poor, since they are based on samples from near present-day sea level. this is also corroborated by munthe (1894, p. 9): “in the south part of the baltic region we possess but comparatively little knowledge of the fauna which results partly from the circumstances that the litorina strata are here to be sought only to a small extent above the sea level”. this has obviously hampered the study within the danish region, so that holocene marine deposits in the baltic are little known up to the present. however, recent activities by marine geologists have given new material from the westernmost part of the baltic – fakse bugt (jensen 1995). as far as the mollusc studies have been submitted (petersen 1994b) but not published in detail, the following comments will be given to fig. 101. the 4 m of sampling from vibrocore 225b comes from the cored section 17.5–13.5 m b.s.l. and has been dated within a time span of 4000 14c years covering the early part of the holocene from the preboreal to the atlantic, through time represented by freshwater, brackish and marine deposits as deduced from the occurrences of mollusc. the loss on ignition shows that the lower third of the sequence has around 25 weight per cent of organic material, while the upper two-thirds of the sequence has less than 5 weight per cent loss on ignition (fig. 101). considering the mollusc species, it appears that the high amount of organic material is not connected solely with the freshwater deposits, but continues into the brackish-water layers. the first rise of sea level is demonstrated by the occurrences of cerastoderma and mytilus and the disappearance of the freshwater molluscs such as valvata macrostoma and the sphaeriidae. the persistent occurrence of the bithynia tentaculata operculae in the oldest part of the brackish-water deposits shows that salinity was lower than 12‰. within this interval, seven mollusc species occur (see fig. 101), viz.: two hydrobia and one rissoa species and the bivalves mytilus, macoma and cerastoderma. the cerastoderma species are rather difficult to identify in all the samples because of their poor state of preservation. at a level of about 15.5 m b.s.l., the change to higher diversity of marine molluscs occurs with such new species as aporrhais pespelicani, nassarius reticulatus and scrobicularia plana besides the steady occurrence of littorina littorea, retusa truncatula and corbula gibba. even bittium reticulatum has been found in one sample. this species tolerates only water with a salinity above 25‰ (sorgenfrei 1958). in all circumstances the more prolific marine fauna is well demonstrated in this part of the sequence, and furthermore, single occurrences of species such as rissoa membranacea, lacuna vincta, and parvicardium exiguum sustain this view. the occurrence of freshwater gastropods is a product of transport from the nearby land, where freshwater streams run into the bay or may be eroded from older deposits by currents. in the topmost part of the core especially fishbones occur up to the present sea bottom at 13.5 m b.s.l. the faunal development in the 4 m core reflects as the oldest element a small lake dated to the time span 9370 ± 135 (k-5649) up to 7900 ± 115 (k-5652) in 14c years b.p. around 7900 b.p. the marine influence is found at the present level of 16.5 m b.s.l. with the establishment of a brackish-water fauna. at a level of 15.30 m b.s.l., which has been dated to 6520 ± 135 14c years b.p. (aar-633), the change to the more prolific littorina sea fauna has taken place. but, the samples geus bulletin no 3.pmd 28-06-2004, 08:46136 137 from higher up in the core do not show any development into the present-day molluscan fauna, first of all because of the absence of mya arenaria that characterises the present-day baltic sea. in conclusion, the littorina sea fauna from the fakse bugt as demonstrated from this core is from the atlantic and shows the molluscs of the transgression around 8000 b.p. (14c years) and the established marine fauna of the atlantic. the fauna from this core contains most of the abovementioned species no longer found in the baltic, and it is shown that this fauna was most probably established during the atlantic. therefore, the marine molluscs from the cored section immigrated to the baltic during the atlantic (appendix 6). among the 15 species with dated appearances in the atlantic only mytilus edulis and macoma balthica have their climatic range within the subarctic–boreal– lusitanian regions, while the rest belong to the boreal– lusitanian group as in the bælt sea area. however, the number of mollusc species from this region is very low with the 19 species found forming only 7.7% of the subfossil fauna. as in the case of the bælt sea, the more prolific fauna can be connected with the atlantic, and this could be caused by the higher tidal amplitude during this time span giving a higher salt content. in the introduction it was mentioned that there is 0 6 12 18 24 30 14 15 16 17 6520 + 135 bp7900 + 115 bp9250 + 130 bp9180 + 130 bp9370 + 135 bpsand gyttja peat lithology li tt or in a lit to re a (l ., 17 58 ) h yd ro bi a ve nt ro sa ( m on ta gu , 1 80 3) h yd ro bi a ul va e (p en na nt , 1 77 7) r is so a al be lla l ov én , 1 84 6 bi tt iu m r et ic ul at um ( da c os ta , 1 77 8) a po rr ha is p es pe lic an i ( l. , 1 75 8) h in ia r et ic ul at a (l ., 17 58 ) r et us a tr un ca tu la ( br ug ui èr e, 1 79 2) m yt ilu s ed ul is l ., 17 58 c er as to de rm a ed ul e (l ., 17 58 ) c er as to de rm a gl au cu m ( po ir et , 1 78 9) m ac om a ba lth ic a (l ., 17 58 ) sc ro bi cu la ri a pl an a (d a c os ta , 1 77 8) c or bu la g ib ba ( o liv i, 17 92 ) 18 –2 0‰ 1 ‰ 10 –2 0‰ 15 –2 0‰ 25 ‰ 18 –2 0‰ 18 –2 0‰ 18 –2 0‰ 6– 10 ‰ 6‰ ~ 6 ‰ 4– 6 ‰ 15 –2 0‰ 15 –2 0‰ ≥ ≥ ≥ ≥ ≥ ≥ ≥ ≥ ≥ ≥ ≥ ≥ ≥ ≥ sp h a er iid a e fo ss il ia v a r ia v al va ta c ri st at a m ül le r, 1 77 4 v al va ta p is ci na lis ( m ül le r, 1 77 4) v al va ta m ac ro st om a st ee nb uc h, 1 84 7 be th yn ia t en ta cu la ta ( l. , 1 75 8) ly m na ea s ta gn al is ( l. , 1 75 8) pl an or bi s co rn eu s (l ., 17 58 ) a ni su s ca ri na tu s (m ül le r, 1 77 4) a ni su s co nt or tu s (l ., 17 58 ) a ni su s cr is ta ( l. , 1 75 8) a ni su s co m pl an at us ( l. , 1 75 8) a ni su s la ev is ( a ld er , 1 83 8) 2‰ 2‰ 0‰ 3– 4 ‰ 5‰ 2‰ 3‰ 5‰ 2– 3 ‰ 4‰ 2– 3 ‰ ≤ ≤ ≤ ≤ ≤ ≤ ≤ ≤ ≤ ≤ ≤ molluscan species with their salinity tolerances loss on ignition weight per cent, %well no, vc225b depth below sea level, m 14c years age the occurrences of molluscan species in the fakse bugt well fig. 101. the occurrences of molluscan species in the fakse bugt well no. vc225b with indication of loss on ignition and sedimentary log with datings (14c years) based on jensen (1995). geus bulletin no 3.pmd 28-06-2004, 08:46137 138 no record of mollusc species from the geological mapping of the island of bornholm (grönwall & milthers 1916). several bottom samples have been analysed from east of bornholm by the author, showing large amounts of tridonta elliptica and tridonta borealis, but the material was never dated, so they could be subrecent specimens. they have therefore not been included in the present list of subfossil holocene molluscs from the baltic. the kattegat climatic regions: asb. class gastropoda subclass opisthobranchia order anaspidea retusa obtusa (montagu 1803) total for climatic regions asb. : 1 (2.2%) climatic regions: asbl class gastropoda subclass prosobranchia order neotaenioglossa littorina saxatilis (olivi 1792) class bivalvia subclass pteriomorpha order mytiloida musculus discors (linnaeus 1767) subclass heterodonta order myoida mya truncata linnaeus 1758 hiatella rugosa (linnaeus 1758) total for climatic regions asbl: 4 (8.9%) climatic regions: .sb. class bivalvia subclass heterodonta order veneroida arctica islandica (linnaeus 1767) total for climatic regions .sb. : 1 (2.2%) climatic regions: .sbl class gastropoda subclass prosobranchia order neotaenioglossa littorina obtusata (linnaeus 1758) lacuna vincta (montagu 1803) skeneopsis planorbis (fabricius 1780) onoba semicostata (montagu 1803) order neogastropoda buccinum undatum linnaeus 1758 subclass heterobranchia order heterostropha omalogyra atomus (phillippi 1841) class bivalvia subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 order pterioida heteranomia squamula (linnaeus 1758) subclass heterodonta order veneroida mysella bidentata (montagu 1803) macoma balthica (linnaeus 1758) total for climatic regions .sbl: 10 (22.2%) climatic regions: ..b. class bivalvia subclass heterodonta order myoida mya arenaria linnaeus 1758 total for climatic regions ..b. : 1 (2.2%) climatic regions: ..bl class gastropoda subclass prosobranchia order archaeogastropoda theodoxus fluviatilis (linnaeus 1758) order neotaenioglossa littorina littorea (linnaeus 1758) littorina tenebrosa (montagu 1803) hydrobia ulvae (pennant 1777) hydrobia ventrosa (montagu 1803) onoba vitrea (montagu 1803) rissoa albella lovén 1846 rissoa inconspicua alder 1844 rissoa membranacea (j. adams 1800) bittium reticulatum (da costa 1778) order heterogastropoda triphora adversa (montagu 1803) order neogastropoda hinia reticulata (linnaeus 1758) subclass heterobranchia order heterostropha brachystomia eulimoides hanley 1844 chrysallida spiralis (montagu 1803) subclass opisthobranchia order anaspidea retusa truncatula (bruguière 1792) akera bullata müller 1776 class bivalvia geus bulletin no 3.pmd 28-06-2004, 08:46138 139 subclass pteriomorpha order pterioida ostrea edulis linnaeus 1758 subclass heterodonta order veneroida acanthocardia echinata (linnaeus 1758) parvicardium exiguum (gmelin 1791) parvicardium scabrum (philippi 1844) cerastoderma edule (linnaeus 1758) cerastoderma glaucum (poiret 1798) scrobicularia plana (da costa 1778) abra alba (wood 1802) paphia aurea (gmelin 1791) tapes decussatus (linnaeus 1758) venerupis pullastra (montagu 1803) order myoida corbula gibba (olivi 1792) total for climatic regions ..bl: 28 (62.2%) the holocene kattegat: 45 (18.2%) the molluscs from the kattegat region have to a great extent been collected in the raised marine forelands to the kattegat itself during the geological mapping in contrast to some of the sampling localities referred to in the preceding two regions. so, although part of the information from djursland comes from borings to a depth of about 10 m b.s.l., it does not present deeperwater deposits as met with in the kattegat proper (petersen 1993). this is also true of the maximum palaeodepth reached, when it is taken into account that the present area is within the isostatic uplift zone with the highest marine shoreline – up to about 10 m – within this region (mertz 1924). forty-five mollusc species have been recorded from the kattegat region. there are no purely lusitanian species, but boreo-lusitanian species. reflecting the above-mentioned facts on maximum palaeodepth, all the species can be found within the tidal/shallow-water zone, viz.: littorina saxatilis, littorina obtusata, littorina tenebrosa, hydrobia ulvae, theodoxus fluviatilis, mytilus edulis, cerastoderma edule, and tapes decussatus. tapes decussatus and paphia aurea are no longer found in danish waters, but occur off western and southern norway in the immediate neighbourhood. rørdam (1891) discussed at some length the different mollusc assemblages in north-eastern sjælland and related these faunas to their relative positions from the open sea and into the innermost part of the fjords. petersen (in rørdam 1891, pp. 106–111), who determined most of the molluscan species in rørdam’s thesis (rørdam 1891, p. 106), stated that: “the deposition of the tapes layers has happened in a period where the danish waters from a hydrographical point of view have been more like the north sea or the open sea than now”. this was clearly demonstrated by the large oyster banks present far into the roskilde fjord in north-eastern sjælland, and it has also been recorded from other mapped areas on fyn and in jylland. the innermost part of the former fjords always carried a rich subfossil mollusc fauna compared to the recent fjord complexes. recent studies from jylland of the faunal changes through time have shown that the rich faunas in the innermost part of such fjord regions developed during the atlantic. this situation is most probably connected with a higher tidal amplitude in the inner danish waters during the atlantic than in the following subboreal period, as argued in petersen (1993). the absence of purely lusitanian species in the midholocene fauna from the kattegat region shows, as mentioned above, that climatic changes have not been of major significance. the most characteristic species for the holocene subfossil fauna in this region – the tapes species – is still to be found within the boreal region. of the 25 dated molluscan species from djursland, 23 species have immigrated during the atlantic along with the transgression (petersen 1993, table 1). only two species, littorina tenebrosa and onoba semicostata, immigrate after the change to more brackish-water conditions that prevailed in the subboreal. among the species with a dated appearance in the atlantic, which form nearly half of the mollusc species known from this region, the dominating climatic group is the boreo-lusitanian, with 18 species. the whole fauna of 45 species from the kattegat region constitutes 18.2% of the subfossil molluscs. this is close to the percentage for the bælt sea area but much less than could be expected regarding the number of species from the recent kattegat region. this recalls the above-mentioned fact that the sampling of the molluscan material has been from raised marine forelands and therefore does not include the deeperwater fauna from the vast area of the kattegat proper. however, as mentioned earlier in connection with the late weichselian marine deposits, new studies are progressing recording the deeper-water fauna with, inter alia, turritella communis from the kattegat during the holocene. geus bulletin no 3.pmd 28-06-2004, 08:46139 140 the limfjord climatic regions: asb. class gastropoda subclass prosobranchia order archaeogastropoda margarites helicinus (phipps 1774) order neogastropoda oenopota turricola (montagu 1803) subclass opisthobranchia order anaspidea retusa obtusa (montagu 1803) class bivalvia subclass heterodonta order veneroida tridonta borealis schumacher 1817 total for climatic regions asb. : 4 (2.7%) climatic regions: asbl class gastropoda subclass prosobranchia order archaeogastropoda acmaea tessulata (müller 1776) order neotaenioglossa littorina saxatilis (olivi 1792) lacuna pallidula (da costa 1778) subclass opisthobranchia order bullomorpha cylichna alba (brown 1827) order anaspidea diaphana minuta brown 1827 class bivalvia subclass palaeotaxodonta order nuculoida nuculoma tenuis (montagu 1808) subclass pteriomorpha order mytiloida musculus discors (linnaeus 1767) subclass heterodonta order veneroida thyasira flexuosa (montagu 1803) order myoida mya truncata linnaeus 1758 hiatella arctica (linnaeus 1758) total for climatic regions asbl: 10 (6.8%) climatic regions: .sb. class bivalvia subclass heterodonta order veneroida arctica islandica (linnaeus 1767) order myoida zirfaea crispata (linnaeus 1758) total for climatic regions .sb. : 2 (1.4%) climatic regions: .sbl class gastropoda subclass prosobranchia order archaeogastropoda acmaea virginea (müller 1776) order neotaenioglossa littorina obtusata (linnaeus 1758) lacuna vincta (montagu 1803) skeneopsis planorbis (fabricius 1780) onoba semicostata (montagu 1803) order neogastropoda nucella lapillus (linnaeus 1758) buccinum undatum linnaeus 1758 subclass heterobranchia order heterostropha omalogyra atomus (phillippi 1841) class bivalvia subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 modiolus modiolus (linnaeus 1758) order pterioida delectopecten vitreus (gmelin 1791) heteranomia squamula (linnaeus 1758) subclass heterodonta order veneroida mysella bidentata (montagu 1803) tellimya ferruginosa (montagu 1803) turtonia minuta (fabricius 1780) parvicardium ovale (sowerby 1840) spisula elliptica (brown 1827) macoma balthica (linnaeus 1758) gari fervensis (gmelin 1791) total for climatic regions .sbl: 19 (12.9%) climatic regions: ..b. class gastropoda subclass heterobranchia order heterostropha chrysallida eximia (jeffreys 1849) total for climatic regions ..b. : 1 (0.7%) climatic regions: ..bl class gastropoda subclass prosobranchia order archaeogastropoda patella vulgata linnaeus 1758 helcion pellucidum (linnaeus 1758) geus bulletin no 3.pmd 28-06-2004, 08:46140 141 iothia fulva (müller 1776) gibbula cineraria (linnaeus 1758) gibbula tumida (montagu 1803) skenea basistriata (jeffreys 1877) order neotaenioglossa littorina littorea (linnaeus 1758) littorina tenebrosa (montagu 1803) lacuna parva (montagu 1803) hydrobia ulvae (pennant 1777) hydrobia ventrosa (montagu 1803) alvania punctura (montagu 1803) cingula semistriata (montagu 1808) onoba vitrea (montagu 1803) rissoa albella lovén 1846 rissoa inconspicua alder 1844 rissoa membranacea (j. adams 1800) rissoa parva (da costa 1779) rissoa violacea desmarest 1814 caecum glabrum (montagu 1803) bittium reticulatum (da costa 1778) turritella communis risso 1826 aporrhais pespelicani (linnaeus 1758) lunatia alderi (forbes 1838) lunatia catena (da costa 1778) order heterogastropoda triphora adversa (montagu 1803) cerithiopsis barleei jeffreys 1867 cerithiopsis tubercularis (montagu 1803) epitonium clathrus (linnaeus 1758) epitonium turtonis (turton 1819) aclis minor (brown 1827) vitreolinaphilippii (rayneval & ponzi 1854) order neogastropoda hinia incrassata (ström 1768) hinia pygmaea (lamarck 1822) hinia reticulata (linnaeus 1758) raphitoma purpurea (montagu 1803) raphitoma linearis (montagu 1803) subclass heterobranchia order heterostropha brachystomia eulimoides hanley 1844 odostomia scalaris macgillivray 1843 chrysallida decussata (montagu 1803) chrysallida indistincta (montagu 1808) chrysallida obtusa (brown 1827) chrysallida spiralis (montagu 1803) ebala nitidissima (montagu 1803) eulimella laevis (brown 1827) eulimella scillae (scacchi 1835) ondina divisa (j. adams 1797) ondina diaphana (jeffreys 1848) odostomia acuta jeffreys 1848 odostomia conoidea winckworth 1932 odostomia turrita hanley 1844 odostomia albella lovén 1846 odostomia plicata (montagu 1803) turbonilla crenata (brown 1827) turbonilla delicata (monterosato 1874) turbonilla lactea (linné 1758) subclass opisthobranchia order bullomorpha acteon tornatilis (linnaeus 1758) cylichna cylindracea (pennant 1777) philine aperta (linnaeus 1767) philine punctata (adams 1800) order anaspidea retusa truncatula (bruguière 1792) retusa umbilicata (montagu 1803) akera bullata müller 1776 class bivalvia subclass palaeotaxodonta order nuculoida nucula nitidosa winckworth 1930 nucula nucleus (linnaeus 1767) subclass pteriomorpha order mytiloida modiolula phaseolina (philippi 1844) modiolus adriaticus (lamarck 1819) modiolaria tumida (hanley 1843) order pterioida aequipecten opercularis (linnaeus 1758) chlamys varia (linnaeus 1758) palliolum striatum (müller 1776) palliolum tigerinum (müller 1776) pododesmus patelliformis (linnaeus 1761) ostrea edulis linnaeus 1758 subclass heterodonta order veneroida lucinoma borealis (linnaeus 1758) lepton nitidum (turton 1822) acanthocardia echinata (linnaeus 1758) parvicardium exiguum (gmelin 1791) parvicardium scabrum (philippi 1844) cerastoderma edule (linnaeus 1758) cerastoderma glaucum (poiret 1798) mactra stultorum (linnaeus 1758) lutraria lutraria (linnaeus 1758) spisula solida (linnaeus 1758) spisula subtruncata (da costa 1778) ensis ensis (linnaeus 1758) phaxas pellucidus (pennant 1777) angulus tenuis (da costa 1778) geus bulletin no 3.pmd 28-06-2004, 08:46141 142 fabulina fabula (gmelin 1791) donax vittatus (da costa 1778) scrobicularia plana (da costa 1778) abra alba (wood 1802) abra nitida (müller 1776) abra prismatica (montagu 1803) chamelea striatula (da costa 1778) clausinella fasciata (da costa 1778) paphia aurea (gmelin 1791) tapes decussatus (linnaeus 1758) timoclea ovata (pennant 1777) venerupis rhomboides (pennant 1777) venerupis pullastra (montagu 1803) mysia undata (pennant 1777) order myoida corbula gibba (olivi 1792) saxicavella jeffreysi winckworth 1930 barnea candida (linnaeus 1758) pholas dactylus linnaeus 1758 subclass anomalodesmata order pholadomyoida thracia phaseolina (lamarck 1818) total for climatic regions ..bl: 107 (72.8%) climatic regions: ...l class gastropoda subclass prosobranchia order archaeogastropoda skenea serpuloides (montagu 1808) order neotaenioglossa alvania lactea (michaud 1830) onoba proxima (forbes & hanley 1850) class bivalvia subclass pteriomorpha order pterioida anomia ephippium linnaeus 1758 total for climatic regions ...l: 4 (2.7%) the holocene limfjord: 147 (59.5%) the mollusc species in the limfjord region amount to 147, and within this relatively high number only four purely lusitanian species are found. skenea serpuloides is known from recent waters in the british isles and southward, alvania lactea and onoba proxima from the western coast of britain and to the south, and anomia ephippium also from the british isles, although including the orkney islands and south to the mediterranean. according to jensen & spärck (1934), anomia ephippium has often not been separated from heteranomia squamula. on the occurrences of anomia ephippium and the three lusitanian gastropods, which are all tiny and difficult species to work with in the subfossil state, the question may then arise whether much reliance should be put in these circumstances indicating that the subfossil holocene fauna of the limfjord region had a more lusitanian affinity than is the case at present. the next climatic group is the boreo-lusitanian, which is by far the largest, with 107 species, 72.8% of all the species recorded from this region. in the limfjord region it is possible to distinguish between the recent fauna that arrived after the breakthrough at the agger tange in 1825, when the limfjord again was established as a saltwater basin, and the older long stage before the middle ages and back to the transgression in the early holocene. the faunal record of the recent mollusc species from the limfjord is an illustration of how fast a population can be established, although not studied from the very beginning (spärck 1943, p. 78). the 85 species established there during the period of about 100 years picture recent immigration (petersen 1986a, p. 223). already the early studies by collin (1884) and petersen (1888) demonstrated that within the subfossil holocene fauna there was a certain number of mollusc species no longer known from the present fauna in the limfjord or within the danish waters at all. these deposits were called the tapes beds by petersen (1888, p. 56). the tapes species sensu petersen (1888) include: paphia aurea, tapes decussatus, and venerupis rhomboides, the last one has lately been recorded as part of the recent danish fauna (jensen & knudsen 1995). the way to have a firmer basis for discriminating between the recent and the holocene fauna in the limfjord region would be to split the boreal climatic region into the three zones: the high-boreal, the midboreal and the low-boreal following the indications on distribution given in the chapter on the molluscan species, following feyling-hanssen (1955) in fig. 4. in this way it turns out that the recent boreo-lusitanian group, with a number of 51, has 45% reaching to the north into the high-boreal, 47% the mid-boreal, and 9% the low-boreal zone. while the subfossil boreolusitanian group of 107 species shows 41% reaching into the high-boreal, 45% the mid-boreal, and 14% the low-boreal zone. there is, so to say, only a slightly higher affinity to more temperate southern waters for the subfossil molluscan fauna, meaning that a hypsothermal period is not clearly demonstrated in the mageus bulletin no 3.pmd 28-06-2004, 08:46142 143 rine environment from the limfjord. in the case of the limfjord region, it should be emphasised that we do have a special situation in connection with the living depths of certain species. taking into consideration again the largest group of molluscs recorded from the holocene limfjord deposits – the boreo-lusitanian group – 12 species have their main occurrence below the tidal zone, and 26 species out of the 107 species encountered occur at a deeper level. when this is seen from the fact that the collection of molluscs during the geological mapping has been done mostly in outcrops, and that the highest marine limit goes up to only about 5 m a.s.l. in this area (mertz 1924), the palaeodepth reached cannot be as deep as figured on the basis of the general depth range of the recent molluscs in danish waters. this was a point stressed by nordmann (jessen 1905, pp. 151–152) and clearly indicates a smaller living depth, found for the molluscs from the limfjord region than from other regions in denmark during the holocene. such a difference in the habitat can still be observed on the beaches of the limfjord, where inter alia the deeper-living mytilid modiolus modiolus is found washed ashore quite commonly in quantities not seen elsewhere along the shores of denmark. the remaining climatic groups are the ones with a distribution extending into the arctic or the subarctic, some with a wide range reaching the lusitanian or the boreal zones to the south. here it should be pointed out that all of the species with their southern limit in the boreal zone have been recorded also from the low-boreal sector, although tridonta borealis is said to be rare in the north sea, but is common in the bælt sea and baltic regions. from this it can be concluded that all the species recorded from the holocene/recent limfjord region could have coexisted in various habitats – except for the three purely lusitanian species. several 14c dates from the limfjord region on molluscs from the holocene reveal the immigration time of a few species (petersen & rasmussen 1995a, table 1), although the actual number of dated shells of certain species are low. from the atlantic: mytilus edulis, cerastoderma edule, ostrea edulis, arctica islandica, spisula subtruncata, corbula gibba, scrobicularia plana, acanthocardia echinata, and venerupis pullastra. from the subboreal: lucinoma borealis, tapes decussatus, and paphia aurea. from the subatlantic: donax vittatus. this is a function of the same species being used in many more datings, because specimens of these species have been present in a sufficient number or weight to allow a conventional 14c dating of the whole assemblage or bed of molluscs. therefore, also the other species present in the samples or stratum shall be considered dated, just as dating of certain levels in borings are taken into account (petersen 1976, 1981, 1985b, 1986c; rasmussen & petersen 1980). in this way a far higher number of first occurrence of species can be demonstrated, still on the basis of absolute dates, as presented in appendix 6. for as many as 114 species out of the total number of recorded species (147) from the holocene mollusc faunas in the limfjord area first occurrence has been dated: 77 species dated to the atlantic, 36 species to the subboreal and one species to the subatlantic. in the atlantic and the subboreal the dominating climatic groups are the boreal–lusitanian species with 55 and 28 species, forming 71.4% and 77.8% respectively. two of the purely lusitanian species which have been dated, alvania lactea and onoba proxima, appeared in the subboreal. the total number of species recorded from the limfjord (147) constitutes 59.5% of the subfossil late quaternary molluscs, which is much higher than seen in the inner danish waters. the north sea climatic regions: asb. class gastropoda subclass opisthobranchia order anaspidea retusa obtusa (montagu 1803) total for climatic regions asb. : 1 (1.1%) climatic regions: asbl class gastropoda subclass prosobranchia order neotaenioglossa littorina saxatilis (olivi 1792) lacuna pallidula (da costa 1778) subclass opisthobranchia order bullomorpha cylichna alba (brown 1827) class bivalvia subclass palaeotaxodonta order nuculoida nuculoma tenuis (montagu 1808) subclass pteriomorpha geus bulletin no 3.pmd 28-06-2004, 08:46143 144 order mytiloida musculus discors (linnaeus 1767) subclass heterodonta order veneroida thyasira flexuosa (montagu 1803) order myoida mya truncata linnaeus 1758 hiatella arctica (linnaeus 1758) total for climatic regions asbl: 8 (8.4%) climatic regions: .sb. class bivalvia subclass heterodonta order veneroida arctica islandica (linnaeus 1767) order myoida zirfaea crispata (linnaeus 1758) total for climatic regions .sb. : 2 (2.1%) climatic regions: .sbl class gastropoda subclass prosobranchia order neotaenioglossa littorina obtusata (linnaeus 1758) lacuna vincta (montagu 1803) order neogastropoda nucella lapillus (linnaeus 1758) buccinum undatum linnaeus 1758 class bivalvia subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 order pterioida heteranomia squamula (linnaeus 1758) subclass heterodonta order veneroida mysella bidentata (montagu 1803) tellimya ferruginosa (montagu 1803) parvicardium ovale (sowerby 1840) spisula elliptica (brown 1827) macoma balthica (linnaeus 1758) total for climatic regions .sbl: 11 (11.6%) climatic regions: ..b. class gastropoda subclass prosobranchia order neotaenioglossa cingula turgida (jeffreys 1870) total for climatic regions ..b. : 1 (1.1%) climatic regions: ..bl class gastropoda subclass prosobranchia order archaeogastropoda gibbula cineraria (linnaeus 1758) theodoxus fluviatilis (linnaeus 1758) order neotaenioglossa littorina littorea (linnaeus 1758) lacuna parva (montagu 1803) hydrobia ulvae (pennant 1777) hydrobia ventrosa (montagu 1803) onoba vitrea (montagu 1803) rissoa albella lovén 1846 rissoa inconspicua alder 1844 rissoa membranacea (j. adams 1800) rissoa violacea desmarest 1814 caecum glabrum (montagu 1803) bittium reticulatum (da costa 1778) turritella communis risso 1826 aporrhais pespelicani (linnaeus 1758) lunatia alderi (forbes 1838) lunatia catena (da costa 1778) order heterogastropoda triphora adversa (montagu 1803) epitonium clathrus (linnaeus 1758) aclis ascaris (turton 1819) aclis minor (brown 1827) aclis walleri jeffreys 1867 order neogastropoda hinia pygmaea (lamarck 1822) hinia reticulata (linnaeus 1758) subclass heterobranchia order heterostropha brachystomia eulimoides hanley 1844 chrysallida indistincta (montagu 1808) chrysallida spiralis (montagu 1803) eulimella laevis (brown 1827) ondina diaphana (jeffreys 1848) odostomia conoidea winckworth 1932 odostomia albella lovén 1846 turbonilla crenata (brown 1827) turbonilla delicata (monterosato 1874) turbonilla lactea (linné 1758) subclass opisthobranchia order bullomorpha acteon tornatilis (linnaeus 1758) order anaspidea retusa truncatula (bruguière 1792) retusa umbilicata (montagu 1803) class bivalvia subclass palaeotaxodonta order nuculoida nucula nitidosa winckworth 1930 geus bulletin no 3.pmd 28-06-2004, 08:46144 145 nucula nucleus (linnaeus 1767) nucula sulcata (bronn 1831) subclass pteriomorpha order pterioida chlamys varia (linnaeus 1758) ostrea edulis linnaeus 1758 subclass heterodonta order veneroida lepton nitidum (turton 1822) acanthocardia echinata (linnaeus 1758) parvicardium exiguum (gmelin 1791) parvicardium scabrum (philippi 1844) parvicardium minimum (philippi 1836) cerastoderma edule (linnaeus 1758) mactra stultorum (linnaeus 1758) spisula solida (linnaeus 1758) spisula subtruncata (da costa 1778) ensis ensis (linnaeus 1758) phaxas pellucidus (pennant 1777) angulus tenuis (da costa 1778) fabulina fabula (gmelin 1791) donax vittatus (da costa 1778) scrobicularia plana (da costa 1778) abra alba (wood 1802) abra nitida (müller 1776) chamelea striatula (da costa 1778) clausinella fasciata (da costa 1778) paphia aurea (gmelin 1791) tapes decussatus (linnaeus 1758) timoclea ovata (pennant 1777) venerupis pullastra (montagu 1803) dosinia exoleta (linnaeus 1758) dosinia lincta (montagu 1803) order myoida corbula gibba (olivi 1792) saxicavella jeffreysi winckworth 1930 barnea candida (linnaeus 1758) subclass anomalodesmata order pholadomyoida cochlodesma praetenue (pulteney 1799) thracia phaseolina (lamarck 1818) total for climatic regions ..bl: 72 (75.8%) the holocene north sea: 95 (38.5%) from the north sea region, 95 species have been recorded, and here no purely lusitanian mollusc species has been found. almost all the north sea finds are recorded also from the limfjord. the boreo-lusitanian group in the north sea region is also, compared to the limfjord, by far the largest group, with 72 species. in this group, only five species can be pointed out as not occurring in the holocene of the limfjord, viz.: parvicardium minimum, clausinella fasciata, dosinia exoleta, dosinia lincta, and cochlodesma praetenue. furthermore, donax vittatus, as remarked earlier in the chapter on the mollusc species, this species does not belong to the limfjord proper, but is recorded from old (subatlantic) beach ridges once facing the skagerrak. today, this bivalve is bound to the exposed coast of denmark, not penetrating into the inner danish waters. the dated occurrences of this species in the north sea area fall in the subatlantic, when the present coastline of jylland was developed (petersen 1994a). donax vittatus has a distribution to the north up into the mid-boreal sector between trondheim fjord and lofoten. the other boreo-lusitanian species in the north sea region which are not in the limfjord go as far north as the high-boreal sector north of lofoten. it is seen that the species not found in the limfjord but in the holocene north sea are to be regarded not as newcomers showing any amelioration, but more probably as representing other conditions prevailing in the north sea area than in the limfjord, as seen in the case of donax vittatus. parvicardium minimum is common only at depths of more than 30 m. clausinella fasciata is far from common in the inner danish waters and, when occurring, is so only at depths of between 15 and 30 m. dosinia exoleta and dosinia lincta can be found from the intertidal zone and out to depths of 70 and 200 m respectively; these species have been recorded from the northern kattegat, although the latter extends into the øresund (jensen & knudsen 1995). the great similarity and the few differences found when the holocene north sea fauna is compared with the limfjord record are explained in the recent study on the agger tange complex (petersen 1994a, 1998). this study, being based on material from several vibrocores west of the agger tange complex, revealed that the jydske rev forms a continuation of the limfjord complex 75 km further towards the west. ams dates of molluscs from this part of the north sea and the oldest cored sections in the agger tange complex show that the marine record can be established from the preboreal and up into the subatlantic (petersen 1985a). the fauna dated from the older part of the holocene is very similar to the limfjord fauna and the younger fauna, containing only a few differences to geus bulletin no 3.pmd 28-06-2004, 08:46145 146 the limfjord fauna, as indicated by the dosinia species. during the older stages of the holocene, the jydske rev complex formed a landscape much the same as that of the present limfjord, while in the younger part of the holocene erosion has taken place and the present coastline of western jylland developed. the appearance of such molluscs as the dosinia and donax species must be connected with the new facies in the exposed coastal areas rather than indicating climatic changes. the well-dated molluscs from the vibrocores in the north sea and borings in the agger tange complex allow the fixing of a first appearance of most of the 95 species recorded from the north sea region as seen in appendix 6 (petersen 1985a, 1994a). the total number of recorded species from the holocene north sea is 95, forming 38.5% of the subfossil molluscs from the late quaternary, which is less than recorded from the limfjord, but out of this number 80 species have been dated with their first appearance. what is of special interest from this area is that the record also covers 26 dated species from the preboreal–boreal and furthermore 27 species from the atlantic, 19 species from the subboreal and eight species from the subatlantic. in all time intervals, the boreo-lusitanian group is by far the dominating part as shown in table 1. as the boreo-lusitanian group in the north sea constitutes the most temperate species, the atlantic can be pointed out as having a slightly higher proportion of warmer mollusc faunal element than the other periods. in order to find the climatic trend for the holocene, the climatic affinity of the recent species not found in the subfossil deposits is compared with that of the subfossil holocene species (table 2). when the affinities to the climatic regions for the 183 holocene subfossil species are compared with those of the 94 recent species that do not occur in the holocene subfossil fauna, a slightly higher affinity to the more temperate regions for the subfossil species appears. from the north sea material it was concluded that the atlantic has a slightly higher proportion of warmer elements than the other periods. so within the danish realm the indication of amelioration, as seen from the molluscan material, points to the atlantic. vendsyssel climatic regions: asb. class gastropoda subclass prosobranchia order neogastropoda oenopota turricola (montagu 1803) subclass opisthobranchia order anaspidea retusa obtusa (montagu 1803) class bivalvia subclass heterodonta order veneroida macoma calcarea (gmelin 1791) total for climatic regions asb. : 3 (2.3%) climatic regions: asbl class gastropoda subclass prosobranchia order neotaenioglossa littorina saxatilis (olivi 1792) lacuna pallidula (da costa 1778) subclass opisthobranchia order anaspidea diaphana minuta brown 1827 class bivalvia subclass pteriomorpha order mytiloida musculus discors (linnaeus 1767) subclass heterodonta order veneroida thyasira flexuosa (montagu 1803) asb. 5 (0.7%) 5 (5.4%) asbl 11 (6.0%) 9 (9.6%) .sb. 3 (1.6%) 0 .sbl 19 (10.4%) 7 (7.4%) ..b. 3 (1.6%) 7 (7.4%) ..bl 136 (74.3%) 64 (68.1%) ...l 6 (3.3%) 2 (2.1%) climatic regions holocene subfossil spp. recent spp. table 2. the climatic trend for the holocene geus bulletin no 3.pmd 09-07-2004, 09:10146 147 order myoida mya truncata linnaeus 1758 hiatella arctica (linnaeus 1758) hiatella rugosa (linnaeus 1758) total for climatic regions asbl: 8 (6.0%) climatic regions: .sb. class bivalvia subclass heterodonta order veneroida arctica islandica (linnaeus 1767) order myoida zirfaea crispata (linnaeus 1758) total for climatic regions .sb. : 2 (1.5%) climatic regions: .sbl class gastropoda subclass prosobranchia order archaeogastropoda acmaea virginea (müller 1776) order neotaenioglossa littorina obtusata (linnaeus 1758) lacuna vincta (montagu 1803) skeneopsis planorbis (fabricius 1780) onoba semicostata (montagu 1803) order neogastropoda nucella lapillus (linnaeus 1758) buccinum undatum linnaeus 1758 subclass heterobranchia order heterostropha omalogyra atomus (phillippi 1841) class bivalvia subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 modiolus modiolus (linnaeus 1758) order pterioida heteranomia squamula (linnaeus 1758) subclass heterodonta order veneroida mysella bidentata (montagu 1803) tellimya ferruginosa (montagu 1803) turtonia minuta (fabricius 1780) parvicardium ovale (sowerby 1840) spisula elliptica (brown 1827) macoma balthica (linnaeus 1758) gari fervensis (gmelin 1791) total for climatic regions .sbl: 18 (13.5%) climatic regions: ..b. class bivalvia subclass heterodonta order myoida mya arenaria linnaeus 1758 total for climatic regions ..b. : 1 (0.8%) climatic regions: ..bl class gastropoda subclass prosobranchia order archaeogastropoda helcion pellucidum (linnaeus 1758) gibbula cineraria (linnaeus 1758) gibbula tumida (montagu 1803) order neotaenioglossa littorina littorea (linnaeus 1758) littorina tenebrosa (montagu 1803) lacuna parva (montagu 1803) hydrobia ulvae (pennant 1777) alvania cimicoides (forbes 1844) alvania punctura (montagu 1803) cingula semistriata (montagu 1808) onoba vitrea (montagu 1803) rissoa albella lovén 1846 rissoa inconspicua alder 1844 rissoa membranacea (j. adams 1800) rissoa parva (da costa 1779) rissoa violacea desmarest 1814 caecum glabrum (montagu 1803) bittium reticulatum (da costa 1778) turritella communis risso 1826 aporrhais pespelicani (linnaeus 1758) lunatia alderi (forbes 1838) lunatia catena (da costa 1778) order heterogastropoda triphora adversa (montagu 1803) epitonium clathrus (linnaeus 1758) epitonium turtonis (turton 1819) vitreolina philippii (rayneval & ponzi 1854) order neogastropoda neptunea antiqua (linnaeus 1758) hinia incrassata (ström 1768) hinia pygmaea (lamarck 1822) hinia reticulata (linnaeus 1758) cytharella coarctata (forbes 1840) raphitoma linearis (montagu 1803) subclass heterobranchia order heterostropha brachystomia eulimoides hanley 1844 odostomia scalaris macgillivray 1843 chrysallida indistincta (montagu 1808) chrysallida obtusa (brown 1827) chrysallida spiralis (montagu 1803) ebala nitidissima (montagu 1803) geus bulletin no 3.pmd 09-07-2004, 09:10147 148 eulimella laevis (brown 1827) odostomia conoidea winckworth 1932 odostomia turrita hanley 1844 odostomia albella lovén 1846 odostomia plicata (montagu 1803) turbonilla lactea (linné 1758) subclass opisthobranchia order bullomorpha acteon tornatilis (linnaeus 1758) cylichna cylindracea (pennant 1777) philine aperta (linnaeus 1767) philine punctata (adams 1800) order anaspidea retusa truncatula (bruguière 1792) retusa umbilicata (montagu 1803) akera bullata müller 1776 class bivalvia subclass palaeotaxodonta order nuculoida nucula nitidosa winckworth 1930 nucula nucleus (linnaeus 1767) subclass pteriomorpha order mytiloida modiolula phaseolina (philippi 1844) modiolus adriaticus (lamarck 1819) modiolaria tumida (hanley 1843) order pterioida aequipecten opercularis (linnaeus 1758) chlamys varia (linnaeus 1758) pecten maximus (linnaeus 1758) pododesmus patelliformis (linnaeus 1761) ostrea edulis linnaeus 1758 subclass heterodonta order veneroida lucinoma borealis (linnaeus 1758) lepton nitidum (turton 1822) kellia suborbicularis (montagu 1803) acanthocardia echinata (linnaeus 1758) parvicardium exiguum (gmelin 1791) parvicardium scabrum (philippi 1844) cerastoderma edule (linnaeus 1758) laevicardium crassum (gmelin 1791) mactra stultorum (linnaeus 1758) lutraria lutraria (linnaeus 1758) spisula solida (linnaeus 1758) spisula subtruncata (da costa 1778) ensis ensis (linnaeus 1758) phaxas pellucidus (pennant 1777) angulus tenuis (da costa 1778) fabulina fabula (gmelin 1791) donax vittatus (da costa 1778) gari depressa (pennant 1777) scrobicularia plana (da costa 1778) abra alba (wood 1802) abra nitida (müller 1776) abra prismatica (montagu 1803) chamelea striatula (da costa 1778) clausinella fasciata (da costa 1778) paphia aurea (gmelin 1791) tapes decussatus (linnaeus 1758) timoclea ovata (pennant 1777) venerupis rhomboides (pennant 1777) venerupis pullastra (montagu 1803) dosinia exoleta (linnaeus 1758) dosinia lincta (montagu 1803) mysia undata (pennant 1777) order myoida corbula gibba (olivi 1792) saxicavella jeffreysi winckworth 1930 barnea candida (linnaeus 1758) pholas dactylus linnaeus 1758 subclass anomalodesmata order pholadomyoida thracia phaseolina (lamarck 1818) total for climatic regions ..bl: 98 (73.7%) climatic regions: ...l class gastropoda subclass prosobranchia order neotaenioglossa alvania lactea (michaud 1830) trivia monacha (da costa 1778) class bivalvia subclass pteriomorpha order pterioida anomia ephippium linnaeus 1758 total for climatic regions ...l: 3 (2.3%) the holocene vendsyssel: 133 (53.8%) in the vendsyssel region the mollusc species amount to 133, nearly the same as recorded from the limfjord region (147). however, some other species occur, although the grouping of species according to the climatic regions to which they belong is almost the same. this is a very important fact considering the large number of species recorded from both regions and according to the conclusions drawn from the limfjord material when compared to the recent fauna in the limfjord, with only a slight difference between the climatic affinities of the subfossil and the recent molgeus bulletin no 3.pmd 09-07-2004, 09:10148 149 luscs, including the recent species which have invaded the limfjord since 1825, but not recorded in the subfossil material. the vendsyssel fauna counts three purely lusitanian elements, but they are far from common, and alvania lactea and trivia monacha are not typical for the former vendsyssel palaeoenvironment. within the boreo-lusitanian group, only ten molluscs are new in the fauna compared to the holocene of the limfjord: alvania cimicoides, neptunea antiqua, cytharella coarctata, pecten maximus, kellia suborbicularis, laevicardium crassum, gari depressa, clausinella fasciata, dosinia exoleta, and dosinia lincta. nordmann (1904) states that the mollusc fauna recorded from the sites north of frederikshavn have a distinct lusitanian affinity. however, as seen above, only ten species are new to the vendsyssel region compared to the limfjord, and here four species reach into the high-boreal sector (north of lofoten), five species into the mid-boreal (between lofoten and trondheim), and only one, neptunea antiqua, has its northern limit within the low-boreal sector (south of trondheim and to the channel). therefore, the designation of the dosinia fauna cannot be one of the special southern appearances. however, nordmann also states that the same species in the dosinia fauna are new compared to the tapes fauna. this is true only for four species, viz. laevicardium crassum, neptunea antiqua, trivia monacha, and kellia suborbicularis, while venerupis rhomboides, lutraria lutraria, gari fervensis and hinia incrassata also occur to the south of frederikshavn and in the limfjord region. furthermore, the characterising species of these beds – dosinia exoleta – is now also recorded from the subatlantic beds at a depth of 32.5 m in the north sea, occurring together with, among others, dosinia lincta. nordmann (1904, pp. 30–31) also argued for the dosinia fauna to be a shallow-water assemblage and discussed in greater detail dosinia exoleta, venerupis rhomboides, lutraria lutraria, laevicardium crassum, lucinoma borealis, arctica islandica, and neptunea antiqua. from the general information at hand on depth relations of these species (chapter on molluscan species), it appears that all except neptunea antiqua can be found from the intertidal zone and out to various depths in deeper water, while neptunea antiqua has a minimum depth of 15 m. taking into account all the recorded molluscan finds from the vendsyssel region, it appears that 84 species (63%) can be found in the tidal–intertidal zone and 48 species (36%) in water deeper than that. therefore it cannot be characteristic of the dosinia fauna that it is a shallow-water assemblage. the new dates which have been used in the discussion of the dosinia fauna date its appearance by the oldest date for the dosinia exoleta, at 4240 b.p. (k5318; in petersen 1991b). skagen climatic regions: asb. class gastropoda subclass prosobranchia order neogastropoda oenopota turricola (montagu 1803) total for climatic regions asb. : 1 (1.4%) climatic regions: asbl class gastropoda subclass prosobranchia order neotaenioglossa lacuna pallidula (da costa 1778) subclass opisthobranchia order bullomorpha cylichna alba (brown 1827) class bivalvia subclass pteriomorpha order mytiloida musculus discors (linnaeus 1767) subclass heterodonta order veneroida thyasira flexuosa (montagu 1803) order myoida hiatella arctica (linnaeus 1758) total for climatic regions asbl: 5 (7.0%) climatic regions: .sb. class bivalvia subclass palaeotaxodonta order nuculoida nuculana minuta (müller 1776) subclass heterodonta order veneroida arctica islandica (linnaeus 1767) total for climatic regions .sb. : 2 (2.8%) climatic regions: .sbl class gastropoda subclass prosobranchia order neogastropoda buccinum undatum linnaeus 1758 geus bulletin no 3.pmd 09-07-2004, 09:10149 150 class bivalvia subclass pteriomorpha order mytiloida mytilus edulis linnaeus 1758 order pterioida heteranomia squamula (linnaeus 1758) subclass heterodonta order veneroida mysella bidentata (montagu 1803) tellimya ferruginosa (montagu 1803) turtonia minuta (fabricius 1780) gari fervensis (gmelin 1791) total for climatic regions .sbl: 7 (9.9%) climatic regions: ..b. class bivalvia subclass heterodonta order myoida mya arenaria linnaeus 1758 total for climatic regions ..b. : 1 (1.4%) climatic regions: ..bl class gastropoda subclass prosobranchia order neotaenioglossa hydrobia ulvae (pennant 1777) barleeia unifasciata (montagu 1803) onoba vitrea (montagu 1803) rissoa albella lovén 1846 rissoa violacea desmarest 1814 bittium reticulatum (da costa 1778) turritella communis risso 1826 aporrhais pespelicani (linnaeus 1758) lunatia alderi (forbes 1838) lunatia montagui (forbes 1838) order heterogastropoda epitonium trevelyanum (johnston 1841) aclis minor (brown 1827) polygireulima sinuosa (sacco 1836) vitreolina philippii (rayneval & ponzi 1854) graphis albida (kanmacher 1798) melanella lubrica (monterosato 1891) melanella alba (da costa 1778) hemiaclis ventrosa (jeffreys ms fricle 1874) order neogastropoda hinia pygmaea (lamarck 1822) hinia reticulata (linnaeus 1758) mangelia brachystoma (philippi 1844) subclass heterobranchia order heterostropha chrysallida decussata (montagu 1803) eulimella scillae (scacchi 1835) odostomia conoidea winckworth 1932 odostomia umbilicaris (malm 1863) turbonilla delicata (monterosato 1874) turbonilla sinuosa (jeffreys 1884) subclass opisthobranchia order anaspidea retusa truncatula (bruguière 1792) retusa umbilicata (montagu 1803) class bivalvia subclass palaeotaxodonta order nuculoida nucula nitidosa winckworth 1930 nucula nucleus (linnaeus 1767) subclass pteriomorpha order pterioida chlamys varia (linnaeus 1758) ostrea edulis linnaeus 1758 subclass heterodonta order veneroida acanthocardia echinata (linnaeus 1758) parvicardium minimum (philippi 1836) mactra stultorum (linnaeus 1758) spisula subtruncata (da costa 1778) phaxas pellucidus (pennant 1777) angulus tenuis (da costa 1778) fabulina fabula (gmelin 1791) tellina pygmaea (lovén 1846) donax vittatus (da costa 1778) abra alba (wood 1802) abra nitida (müller 1776) abra prismatica (montagu 1803) chamelea striatula (da costa 1778) timoclea ovata (pennant 1777) order myoida corbula gibba (olivi 1792) saxicavella jeffreysi winckworth 1930 barnea candida (linnaeus 1758) pholas dactylus linnaeus 1758 subclass anomalodesmata order pholadomyoida lyonsia norvegica (gmelin 1791) cochlodesma praetenue (pulteney 1799) thracia phaseolina (lamarck 1818) total for climatic regions ..bl: 54 (76.1%) climatic regions: ...l class gastropoda subclass prosobranchia order heterogastropoda vitreolina collensi (sykes 1903) geus bulletin no 3.pmd 09-07-2004, 09:10150 151 total for climatic regions ...l: 1 (1.4%) the holocene skagen: 71 (28.7%) from the skagen boring, 71 holocene species have been recorded. the boreo-lusitanian group dominates with 54 species (76.1%) of the holocene mollusc species from the skagen well. one purely lusitanian species, vitreolina collensi, has been found, while all the other species occur in the boreal and to some extent the arctic. in this way the skagen well material resembles that of other regions like the vendsyssel, limfjord and north sea during the holocene. the skagen well material has all been recorded to certain stratigraphical levels, as seen in appendix 6, so the climatic indications through time appear, but the number of molluscan species is very low. in the preboreal/boreal, only three species have been recorded. this has been explained as a result of a deeper water where the echinoids dominate. higher up in the sequence, the number of molluscs increases – 23 species in the subboreal and 68 species in the subatlantic. through the chronostratigraphical levels, the climatic regions of the boreo-lusitanian from the dominating one, and the purely lusitanian vitreolina collensi as mentioned above occurs in the subboreal and subatlantic. however, as already stated, the development of the facies in the skagen well during the holocene does change the environment from the deeper-water facies with few molluscs through the bottom community with turritella communis into the prolific shallow-water community. in this way the youngest part covering the subboreal–subatlantic is also by far the part with the highest species diversity. the environmental changes within the seven regions through the late quaternary evaluated by the molluscan communities met with in the seven stages the seven chronological stages which have been described according to their climatic affinities are seen in fig. 102 and fig. 103 covering the eemian, the early/ middleweichselian, thelateweichselian, thepreboreal/ boreal, the atlantic, the subboreal, and the subatlantic. in this way the climatic cycle during the late quaternary is demonstrated on the basis of marine mollusc species which indicate that the eemian has by far the highest amount of the more temperate species, while the holocene reached its maximum during the atlantic, although only slightly more than the other stages within the holocene, as already commented upon in the previous chapter. it is generally accepted that the eemian summer temperatures were higher – about 2°c above the present. with glaciers smaller than the present day, this means that the sea level was 4–6 m higher than today (andersen & borns 1994, pp. 44–49). and as pointed out by donner (1995, p. 39): “the submergence was clearly greater after the saalian glaciation than after the weichselian and possibly after the older glaciations”. donner sees this in northern europe as “a result of a comparatively great downwarping of the earth’s crust during the extensive saalian glaciation”. the rebound since the last glaciation has come to an end within the danish area (petersen 1985c, 1991b). this means that the eemian deposits, when found in denmark in nonglacio-dislocated state, can be regarded as being in the original position related to sea level, although there might be some movements in relation to neo-tectonic activities, as mentioned earlier. in the light of the observations mentioned above, the seven regions will be discussed according to the environmental characteristics such as the climatic affinities for the molluscs recorded in appendix 6 for each region, as appearing in fig. 103. however, for the holocene still as many as 130 species including the recent ones (95) not found as subfossil have not been dated to give their first appearance, see fig. 102: unknown arrival in holocene. at the end of each of the seven stages the molluscan communities sensu c.g.j. petersen will be presented in tables 3–9. geus bulletin no 3.pmd 09-07-2004, 09:10151 152 eemian stage 130 000 – 115 000 b.p. the bælt sea, region 1 appendix 6 and fig. 103 already forchhammer (1842, p. 64) designated cyprina (arctica) islandica to be the characteristic bivalve of the bælt sea eemian, as known to the present-day geologists. furthermore, forchhammer points out that the characteristic bivalve, cyprina islandica, occurs every where in large quantities, but always in crust specimens. however, all the shell fragments occurring together show that the specimen has been broken after deposition in the clay, most probably by the cataclysms which have given the beds their tilt. johnstrup (1882a, p. 55) points to the indications of the molluscs as being a deposition of a shallow-water sea and also mentioned the mytilus beds. johnstrup points out (1882a, p. 56) that the dislocated floes – as already noticed by forchhammer – have the original succession within each floe, saying that the cyprina clay and the mytilus beds have not been disconnected during the dislocations. later investigations by nordmann (harder 1900; nordmann 1908, 1913) demonstrated that the venus aurea as observed by johnstrup (1882a, p. 66) could be regarded in parts as the no longer living tapes aureus gm. var eemiensis nordmann or tapes senescens doederlein; in this book paphia aurea senescens. the tapes species do represent shallow-water environments (see the chapter on the molluscan species), and therefore the whole of the bælt sea region can be characterised by the three bivalves mentioned above, from the eulittoral to the infralittoral shallow-water zones: mytilus, tapes and cyprina. among the three species mentioned, the tapes species (paphia aurea senescens) also remains as the only subfossil bivalve from denmark which can be regarded as an index fossil from the marine eemian. the stratigraphical position of the marine eemian is according to jessen & milthers (1928, p. 179) contemporaneous with the mixed oak forest zone and the carpinus zone in the interglacial bogs; furthermore, jessen & milthers conclude (1928, p. 341) that the climate of jylland and nw germany in that part of the interglacial period which answers to zone f [culmination of the curves for mixed oak forest] was no less atlantic in character than (100%) (100%) (100%) (100%) (100%) (100%) (100%) 3 11 48 48 16 53 3 290 14 140 36 41 27 93 77 72 486 1 2 6 5 10 19 14 1 1 2 1 11 7 11 2 6 5 6 3 2 4 2 3 2 15 1 5 3 14 7 8 1 1 1 90 21 70 56 53 10 3 1 (3%) (5%) (17%) (12%) (7%) (53%) (34%) (4%) (1%) (3%) (1%) (8%) (19%) (27%) (7%) (6%) (6%) (8%) (2%) (6%) (10%) (2%) (4%) (3%) (11%) (3%) (12%) (11%) (15%) (9%) (11%) (1%) (1%) (1%) (64%) (78%) (75%) (73%) (74%) (7%) (4%) (1%) frequency: number of molluscs column per cent, % climatic affinity eemian early/middle late weichselianweichselian preboreal/ boreal atlantic subboreal subatlantic total number of species unknown arrival in holocene arctic arctic, subarctic arctic, subartic, boreal, lusitanian arctic, subarctic, boreal subarctic, boreal subarctic, boreal, lusitanian boreal boreal, lusitanian lusitanian total number of species column per cent, % 35 s ub fo ss il sp ec ie s 95 r ec en t sp ec ie s pleistocene holocene time: the seven stages fig. 102. the seven stages from the eemian through the weichselian – comprising the early/middle weichselian and late weichselian – to the end of the holocene. preboreal–boreal, atlantic, subboreal, and subatlantic are here shown according to their climatic affinities based on the molluscan record. geus bulletin no 3.pmd 09-07-2004, 09:10152 153 (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (2%) (3%) (10%) (73%) (12%) (2%) (10%) (7%) (9%) (28%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (100%) (5%) (5%) (5%) (21%) (63%) (2%) (2%) (5%) (5%) (3%) (3%) (9%) (72%) (16%) (2%) (4%) (5%) (20%) (1%) (5%) (1%) (11%) (73%) (9%) (2%) (8%) (5%) (12%) (14%) (32%) (16%) (13%) (2%) (11%) (2%) (56%) (16%) (11%) (5%) (7%) (33%) (7%) (7%) (29%) (14%) (43%) (7%) (2%) (6%) (2%) (1%) (4%) (14%) (14%) (29%) (43%) (17%) (8%) (4%) (5%) (3%) (17%) (53%) (19%) (6%) (3%) (17%) (50%) (33%) (11%) (10%) (1%) (25%) (75%) (17%) (5%) (6%) (11%) (37%) (23%) (9%) (14%) (33%) (33%) (23%) (13%) (15%) (6%) (17%) (8%) (5%) (5%) (5%) (4%) (8%) (12%) (77%) (2%) (3%) (4%) (4%) (33%) (67%) (1%) (0%) (6%) (3%) (16%) (74%) (3%) (5%) (6%) (5%) (13%) (87%) (2%) (3%) (22%) (78%) (6%) (4%) (1%) (8%) (3%) (17%) (71%) (2%) (9%) (10%) (15%) (12%) (7%) (7%) (85%) (3%) (2%) (5%) (50%) (50%) (5%) (0%) (100%) (2%) (50%) (50%) (1%) (0%) (6%) (6%) (6%) (78%) (6%) (4%) (3%) (2%) (6%) (8%) (16%) (11%) (5%) (5%) (63%) (5%) (10%) (1%) (33%) (3%) (33%) (67%) (1%) (0%) (4%) (4%) (9%) (78%) (4%) (2%) (5%) (2%) (4%) (4%) (100%) (0%) (13%) (25%) (63%) (2%) (2%) (1%) (1%) (7%) (3%) (10%) (1%) (75%) (1%) (2%) (8%) (10%) (8%) (33%) (11%) (4%) (11%) (11%) (33%) (33%) (11%) 59 19 32 91 55 14 7 36 4 35 9 26 3 31 15 23 77 27 2 8 2 36 19 3 23 1 8 68 7336 12 57 64 20 84 3 462 25 1 1 2 18 1 1 1 2 5 1 5 2 7 1 51 1 8 1 1 2 2 2 28 2 3 2 1 1 12 1 2 1 2 2 1 5 23 2 13 5 18 1 6 2 13 55 2 2 23 1 1 1 1 2 3 1 6 19 7 2 1 1 3 2 4 13 8 3 5 1 1 3 3 1 1 2 3 20 1 2 6 43 7 1 1 1 4 12 1 3 23 5 1 5 1 10 66 8 9 7 1 6 1 31 1 4 2 6 1 age ep oc h pl ei st oc en e h ol oc en e region 1. bælt sea 2. baltic sea 3. kattegat 5. north sea 6. vendsyssel 7. skagen 3. kattegat 6. vendsyssel 7. skagen 6. vendsyssel 7. skagen 5. north sea 7. skagen 1. bælt sea 2. baltic sea 3. kattegat 4. limfjorden 5. north sea 6. vendsyssel 7. skagen 3. kattegat 4. limfjorden 5. north sea 6. vendsyssel 7. skagen 4. limfjorden 5. north sea 7. skagen eemian early/middle weichselian late weichselian preboreal/ boreal atlantic subboreal subatlantic total of species in number row per cent, % column per cent, % total of species: in number row per cent, % column per cent, % arctic a, s boreal b, l lusitanian a, s, b a: arctic, s: subarctic, b: boreal, l: lusitanian a, s, b, l s, b s, b, l climatic affinity fig. 103. climatic characterisation (affinity) of mollusc assemblages for each region over the seven stages (ages) since the eemian. geus bulletin no 3.pmd 09-07-2004, 09:10153 154 the climate of the litorina period in postglacial time. also s.t. andersen (1965, pp. 499–500) points to the eemian as having an oceanic and warm climate. in his diagram from hollerup, zone 5 represents the climax forest. jessen & milthers (1928, p. 179) bring forth the view of nordmann saying that the lusitanian mollusc fauna was moving into the baltic basin by way of the sounds that cut diagonally through the southern portion of the cimbrian peninsula. however, this idea is, as discussed earlier, not accepted by the present author, where a connection to the west is advocated to be over the kattegat–vendsyssel–skagen regions. an eastern open connection to the white sea over finland and russia will be discussed in connection with the baltic region. the baltic, region 2 appendix 6 and fig. 103 in the baltic sea part of denmark the number of recorded marine molluscs from the eemian has fallen to 19 species compared to the bælt sea region, and no purely lusitanian species occur. however, as pointed out earlier, this change in the climatic situation regarding the baltic fauna can be seen as a consequence of the fact that here only the deeper-water assemblage characterised by turritella communis occurs, although one of the species known from the characteristic part of the bælt sea fauna is recorded, i.e. arctica islandica. however, the tapes species are not met with in this region. the deeper-water environment is therefore well characterised by the turritella community. although the decline in number of species very much resembles the present-day situation between the bælt sea and the baltic as mentioned earlier, five species in the eemian fauna, including turritella communis, show a salinity above the present conditions in the baltic. from the study of diatoms at ollala in eastern fennoscandia, forsström et al. (1988, p. 322) write: “this mixture of warm and cold indicators probably means that the eemian sea in the baltic basin had a connection both to the north sea in the west and to the arcticocean via thewhite seabasin in thenortheast”. only a few works from eastern fennoscandia have been based on molluscan studies. however, among the papers by zans (1936), sokolova et al. (1972), and gross (1967), gross mentions the following molluscs: portlandia arctica, clinocardium ciliatum, heteranomia squamula, macoma calcarea, littorina littorea, and cerastoderma edule. here the three last mentioned species occur in the eemian from the danish baltic and clinocardium ciliatum from the vendsyssel region, although here in the upper turritella terebra zone correlated to the early weichselian, as discussed later. portlandia arctica has not been recorded from the danish eemian, although it occurs in the beds below the eemian in the anholt boring (the kattegat region), where seidenkrantz (1993, p. 284) also has demonstrated foraminiferal zones a–d with arctic species. gross (1967, p. 118) regards the arctic and arctic– boreal molluscs in the eemian clay as: “relikte aus der portlandia-transgression des dnepr ii-spätglazials, die nach dem pollen-profil und -diagramm der eem-transgression voranging”. the older correlation of the so-called weissmeer transgression by zans (1936, table 1) contains further details on the molluscs upon which the correlation has been based, and it also includes the danish area, mentioning ‘dänische inseln’ after ødum (1933) and ‘skaerumhede’ after jessen et al. (1910). however, the occurrence of the high-arctic portlandia arctica should be placed in the late saalian, as seen in the kattegat region mentioned above, while the arctic–boreal species clinocardium ciliatum could be taken as a relict in the skærumhede sequence from the late saalian environment within the danish area or introduced by the cooling in the early weichselian. the faunal development has been worked out in more detail between the eastern fennoscandia and the danish area (funder et al. 2002). the connection to the arctic over the white sea during the eemian seems to be well established, but only for a shorter time, 1000–2000 years of the more than 10000 years that the eemian sea existed in the baltic region (funder 2000, p. 68). the kattegat, region 3 appendix 6 and fig. 103 from the kattegat region, 32 species have been recorded, with a high amount of boreo-lusitanian species (72%). five species are lusitanian, among which are found the characteristic eemian species of the shallow-water environment, including the tapes species. however, also the deeper-water environment is represented by the turritella community in this region. geus bulletin no 3.pmd 09-07-2004, 09:10154 155 on anholt, the turritella community occurs at a depth of around 70 m b.s.l., and the tapes fauna in the isefjord area at ejby at a depth of around 10 m a.s.l. the latter is considered to be in situ (madsen 1968). it is tempting to regard the two localities as being at about their original elevation in relation to an eemian sea level some what higher than the recent one, since the glacio-isostatic rebound had expired (petersen 1991b). on the basis of the scattered eemian localities of which some are floes in the weichselian glacial deposits, the maximum extent of the eemian sea cannot be given. however, both the shallow-water environment characterised by the tapes species and the deeper water by the turritella species have been demonstrated. in this way both of the characteristic marine environments from the bælt sea and the baltic respectively are represented in the kattegat region. the north sea, region 5 appendix 6 and fig. 103 the largest amount of mollusc species within the eemian have been recorded from the north sea region, or to be more precise from the coastal region of the north sea. in the danish part of the north sea, many studies on microfossils from the oil and gas fields have demonstrated eemian deposits in the central north sea, but their macrofossils have not been studied (knudsen 1985a, 1986). however, the large amount (91) of molluscs from the coastal region fall into different facies, as seen in the previous regions when a much lower number was looked at. gripp (1964) uses the ‘senescens sand’ and ‘turritella ton’ to give his idea of the marine order of the strata. however, in this context, working with regions and not with localities, it should be emphasised that the development of different facies most probably happened in parallel. gripp (1964, p 223) expresses this himself in saying: “tapes-sand und turritellen-ton sind die beiden facies, die während des ansteigs des meeresspiegels entstanden”. as seen from the species found in the danish north sea coastal region, we do find the turritella species and tapes species, but also the occurrences of donax vittatus should be mentioned as a facies indicator, characterising the high-energy coastal environment from this area facing the eemian north sea. from recent studies on foraminifera in northern germany at the kiel canal, knudsen (1986) shows that the marine transgression took place in the warm part of the eemian, and hinsch (1985) in his mollusc study from the same area revealed three mollusc communities characterising the shallow-water environment, with such genera as mytilus-cerastoderma, acanthocardiavenerupis and bittium-varicorbula. however, the old material from many localities in the danish north sea region cannot be worked out to such detail, although all the marine mollusc species mentioned by hinsch (1985) have been recorded from the danish eemian north sea region. when the molluscan fauna in the danish eemian north sea region is compared with the eemian on the west coast of norway as described by mangerud et al. (1981), 20 molluscan species out of the 35 species recorded from the fjøsangerian are known from the north sea region and 7 species from other eemian regions in denmark. here macoma calcarea and nuculana pernula belonging to the arctic–boreal group occur in the vendsyssel and skagen regions, and in the baltic region macoma calcarea representing deeper water during the eemian, while chlamys islandicus, which does not occur in the danish eemian deposits, is found in the weichselian recorded from the vendsyssel region. this means that the fjøsangerian can be regarded as slightly cooler than the danish north sea eemian deposits. another marine eemian deposit in norway described by andersen et al. (1983) at bø on karmøy (sw norway) revealed 25 molluscan species from the avaldsnes interglacial described in detail by sejrup (1987). here as many as 20 species are in common with those in the danish north sea region, and one, hinia incrassata, has been recorded from the eemian in the vendsyssel region. four species have not been found in the danish eemian. these are the arctic–lusitanian and arctic– boreal species puncturella noachina and boreotrophon clathratus respectively, the latter occurring in the late weichselian deposits in the vendsyssel region. this is much in line with the observations from the fjøsangerian deposits. the four non-occurrences in the danish eemian among the boreo-lusitanian species pecten maximum and lucinoma borealis, the latter being common at both norwegian localities, are difficult to explain. among the 8 purely lusitanian species recorded from the north sea region, only plagiocardium papillosum occurs in the norwegian eemian at fjøsanger, which again points to a slightly cooler position for the norwegian localities. geus bulletin no 3.pmd 09-07-2004, 09:10155 156 the vendsyssel region, region 6 appendix 6 and fig. 103 the eemian mollusc fauna from vendsyssel comprises 55 species with no purely lusitanian climatic affinity, while quite a few (nine species~ 16%) are found in the arctic and the boreal zones. the stratigraphical position has been well elucidated through foraminiferal investigations (knudsen & lykke-andersen 1982; knudsen 1984, 1985b, 1992; lykke-andersen 1987). the study by lykke-andersen (1987, fig. 5) also involves the molluscs, and references are made to the zones established on the basis of macrofossils. according to the foraminiferal studies, the transition to the early weichselian takes place around 120 m b.s.l. in the skærumhede i sequence (jessen et al. 1910), which is about 140 m b.s. this means that the upper part of the turritella terebra zone falls within the early weichselian. the two arctic to high boreal species serripes groenlandicus and clinocardium ciliatum at depths of 132 m and 127 m b.s. respectively are discussed by nordmann (jessen et al. 1910, pp. 124–128), and the climatic indications from turritella communis mean that the assemblage existed at the transition between the high and middle boreal. in the paper by knudsen (1992), it is said that an abrupt faunal change at the eemian–weichselian boundary reflects a drop in water depth of at least 50 m and a subsequent drop in temperature of several degrees. the drop in temperature might well be reflected in the mollusc fauna by the occurrences of the two bivalves mentioned above, and for the drop in sea level it is tempting to recall the observed occurrence of the eulittoral mytilus edulis at a depth of 135 m b.s.l. in the skærumhede i boring and up to the abra nitida zone, which forms the transition to the arctic turritella erosa community as mentioned earlier. therefore, within the vendsyssel area the eemian (isotopic stage 5e) is represented by a turritella community that continues into the beds representing the isotopic stages 5d–a (knudsen 1992, fig. 4). the hordalandian stage in western norway contains serripes groenlandicus and clinocardium ciliatum species and is referred to the early weichselian (mangerud et al. 1981). arctic conditions first occurred in the macrofossil zones turritella erosa, balanus crenatus and macoma calcarea (bahnson et al. 1974), which cover the portlandia arctica zone sensu nordmann (jessen et al. 1910, fig. 8). the skagen region, region 7 appendix 6 and fig. 103 the 14 eemian molluscs found in the skagen boring are the lowest number recorded within the eemian sites. however, the finds are a clear omen of the deeperwater environment not encountered earlier in denmark on the basis of molluscs. the recorded molluscs point to an environment like the deeper part of the skagerrak today, with a community such as the amphilepis norvegica/pecten vitreus, where the latter (delectopecten vitreus) occurs in the skagen well, as mentioned earlier. the boundary to the overlying arctic deposits characterised by the occurrences of portlandia arctica is sharp and coincides with a sedimentological change to a diamicton with dropstones in the arctic part, as found in the skærumhede sequence within the turritella erosa zone (bahnson et al. 1974). therefore, in the skagen well no transition zone from substage 5e to 5d–a can be demonstrated in the molluscan faunas. the recorded eemian communities and/or characteristic molluscan species for six regions with eemian marine deposits are given in table 3. 1. bælt sea mytilus edulis littoral tapes spp. shallow arctica islandica deeper 2. baltic littoral shallow turritella t. communis deeper 3. kattegat mytilus edulis littoral tapes spp. shallow turritella t. communis deeper 5. north sea donax vittatus littoral tapes spp. shallow turritella t. communis deeper 6. vendsyssel littoral shallow turritella t. communis ~ 100 m 7. skagen littoral shallow amphilopsis/pecten delectopecten vitreus > 100 m region community species depth table 3. eemian communities and/or characteristic molluscan species geus bulletin no 3.pmd 09-07-2004, 09:10156 157 early/middle weichselian stage 115 000 – 25 000 b.p. the kattegat, region 3 appendix 6 and fig. 103 although the number of mollusc species recorded from the kattegat region during the early/middle weichselian is low – seven species – the climatic indications for the arctic environment are clear, considering that all species canbe found in thehigh arctic, and that one species, portlandia arctica, is high arctic par excellence, and macoma calcarea indicates shallow water. the stratigraphical position of the arctic macoma community found in the kattegat region – holmstrup on sjælland – has been determined by foraminiferal correlation and aminostratigraphical investigations, as mentioned earlier (petersen & buch 1974; miller & mangerud 1985). recently, the foraminiferal studies of the quaternary sequence in the anholt boring have demonstrated a middle weichselian deposit at a depth of about 50 m b.s.l. (seidenkrantz 1993). considering the information given by knudsen (1992) on a drop of sea level of around 50 m during the transition from the eemian to the weichselian, the middle weichselian beds in the cored section on anholt may represent rather shallow-water deposits. this is in accordance with the occurrence of the arctic macoma community. the vendsyssel region, region 6 appendix 6 and fig. 103 foraminiferal studies by lykke-andersen (1987) indicate that the early weichselian beds are represented by the upper turritella terebra zone, and that the deeper-water temperate turritella community continued into the first part of the weichselian with the following abra nitida zone as a transition to the arctic deeper-water turritella community (turritella erosa). together with the arctic and arctic–subarctic species (seven in number, forming nearly 20% of the mollusc species), the sedimentological data show the occurrences of ice-rafted material recorded both from the skærumhede i and the skærumhede ii borings (jessen et al. 1910, p. 76; bahnson et al. 1974, figs 3, 4, 7) reflecting arctic conditions. it has been argued by nordmann (jessen et al. 1910) that species within the genera mytilus, cyprina, zirphaea, nassa, and bittium must be regarded as allochthonous and older elements. however, they could also be regarded as stray finds from contemporary shallow-water to littoral deposits occurring within times of higher temperatures in the near-shore areas, similar to the near-shore fauna of middle weichselian age – the bø interstadial (40–64 ka) with gibbula cineraria and mytilus edulis, the latter occurring frequently (sejrup 1987). during the younger part of the marine middle weichselian – around 32 000 b.p. – when the shallow-water arctic macoma community was established, no mytilus edulis or bittium reticulatum have been recorded. the development of the bottom communities within the older weichselian sequence is therefore given by the transition from the turritella communities in deeper water to the arctic macoma community in shallow water. the skagen region, region 7 appendix 6 and fig. 103 very few molluscs have been found in the skagen boring of older weichselian age. the four species are all arctic, and the occurrence of portlandia arctica shows that high arctic conditions have prevailed and ice-rafted material occurs. there is no indication of near-shore fauna as recorded from the vendsyssel region. from this, it might be concluded that the deposition of these beds took place in the first part of the middle weichselian, contemporaneous with the deposition of the turritella erosa beds of the skærumhede sequence, but at a water depth of more than 100 m, as demonstrated earlier. from the few finds, it is not safe to point to a certain community on the basis of molluscs. however, the community in deeper arctic waters is described by other animals than molluscs, i.e. the ascidia–spongia epifauna, and at depths exceeding 200 m by gorgonocephalus species. from the estimate on water depth taken in comparison with the early part of the arctic sequence in the vendsyssel region, the palaeodepth must have been well above 100 m. the recorded early/middle weichselian communities and/or characteristic molluscan species are given in table 4. geus bulletin no 3.pmd 09-07-2004, 09:10157 158 late weichselian stage 25 000 – 10 000 b.p. the vendsyssel region, region 6 appendix 6 and fig. 103 the 35 mollusc species found in the vendsyssel region are in number very close to the number of species encountered in the vendsyssel region during the early/ middle weichselian (36 species). however, as seen in fig. 103, the percent of molluscs with a wide range and only connected to the subarctic and southwards is higher in the late weichselian (46%) compared to the early/middle weichselian (28%). this can be explained by the way the development in the two seas before and after the main glaciation, the older and younger yoldia sea respectively took place. the deposits from the older yoldia sea reflect the transition from deeper arctic to shallow-water arctic communities, the turritella and macoma communities respectively. the late weichselian beds within the shallow-water environment show a development from the arctic macoma community to the boreo-arctic mytilus-zirphaea community after 13 000 b.p., with a deeper-water community characterised by the portlandia arctica species, as outlined by petersen (1984), which could be part of the deeper macoma community – the so-called ophiocten zone. these observations form the background for the earlier given explanation of the occurrences of boreal shallow-water species such as mytilus edulis in the deeper-water arctic community in the older weichselian deposits from the vendsyssel region. the occurrences of the mytilus edulis species in the late weichselian deposits in large quantities are described by jessen (1899). the dates of the earliest occurrences of zirphaea and mytilus go back to 12 770 and 12 520 b.p. 14c years respectively. all the 30 14c dates forming the base for the evaluation of the late weichselian sea levels and occurrences of fauna communities as figured in petersen (1984, fig. 1) have been listed by petersen & rasmussen (1995a, table 1). it appears that the dates older than 13 000 b.p. 14c years all come from hiatella arctica and macoma calcarea (only one date) going as far back as 14 650 ± 190 b.p. 14c age. considering the mollusc species inhabiting the waters “of the swedish west coast shortly after deglaciation” (fredén 1986, p. 55), one finds also chlamys islandica and mytilus edulis shortly after 13 000. the latter within the time span of its first dated occurrence in the younger yoldia sea deposits in denmark. as to the deeper-water deposits also around 13 000, one can take the sample from the dybvad clay pit (fig. 1) dated to 13 010 ± 190 b.p. 14c which contained the following molluscs: hiatella arctica, mya truncata, macoma torelli, portlandia arctica, buccinum groenlandicum and cylichna occulta (petersen 1984). it is seen that still after the immigration of boreo-arctic fauna to the shallow-water environment, showing an amelioration in climate, the arctic community persisted in the deeper water as shown by the dating from the bindslev clay pit giving 12 650 ± 180 14c age b.p., with species such as macoma calcarea and portlandia arctica. the same situation can be estimated from the investigation of the pleistocene/holocene boundary in south-western sweden (the moltemyr core) where “zone z comprising the samples from 560 cm to 650 cm, is characterised by portlandia arctica and nucula tenuis (nuculoma tenuis, here taken as a species with a wide climatic range), and by the absence of many of the species of the overlaying zone (such as mytilus edulis) … the water depth during deposition of zone z was greater than during any of the other zones (above) probably more than 20 m” (feyling-hanssen 1982, p. 128). regarding the climatic indication of portlandia arctica, feyling-hanssen (1982, p. 131) quotes andersen (1975, p. 54) saying: “evidently, portlandia arctica lived near the ice fronts [in southern norway] also during older, glacial phases, but it seems to have disappeared from our coasts shortly after the ra event, probably due to a warming of the sea”. 3. kattegat littoral arctic macoma m. calcarea shallow deeper 6.vendsyssel littoral arctic macoma m. calcarea shallow arctic turritella t. erosa ~ 90 m turritella t. communis ~ 90 m 7. skagen littoral shallow ascidia–spongia 100–200 m region community species depth table 4. early/middle weichselian communities and/or characteristic molluscan species geus bulletin no 3.pmd 09-07-2004, 09:10158 159 according to sørensen (1979), the disappearance of portlandia arctica from the oslo fjord area occurred somewhat before 10 000 b.p. the younger dryas marine deposits have not been demonstrated in the vendsyssel region but are recorded from the skagen region, which will be discussed next. the skagen region, region 7 appendix 6 and fig. 103 the nine species recorded from the skagen well indicate an arctic deeper-water community very much like the arca-astarte crenata community (ockelmann 1958). there are no finds of species which could be referred to the more shallow-water environment as seen in the case of the vendsyssel region to the south both during the early/middle weichselian and the late weichselian. through most of the history of the cored section of pleistocene age, the skagen well has revealed mollusc assemblages from a deeper-water environment. also the transition to the holocene takes place in deeper water. the purely arctic species portlandia arctica together with bathyarca glacialis is found right up to the strata dated to around 10 000 b.p. forming the pleistocene– holocene boundary in the skagen well. however, in the description of the shell fauna of the marine clays in the oslo fjord region, brögger (1900, p. 685) states: “portlandia arctica is never found in the arca clay”. later datings of the ‘middle arca clay’ and the ‘younger arca clay’ given by b.g. andersen(1965,p.118) yielded early preboreal ages. in the younger arca clay from norway, species such as mytilus edulis, zirfaea crispata, and macoma balthica (b.g. andersen 1965, table 2) are also found, which characterise the shallowwater deposits in vendsyssel after 13 000 b.p. 14c age. the recorded recent occurrence of bathyarca glacialis from southern iceland implicates extension into the high-boreal region, although the main extension is in the arctic. the norwegian records of bathyarca glacialis come from a more shallow-water environment, as seen from the occurrences of the three shallowwater species mentioned above. the recorded late weichselian communities and/or characteristic molluscan species are given in table 5. the preboreal–boreal stage 10 000 – 8000 14c years b.p. the north sea, region 5 appendix 6 and fig. 103 mollusc faunas from the late weichselian have here been recorded only from the vendsyssel and skagen regions. as seen in fig. 103, the early part of the holocene, the preboreal and boreal, have a record of 26 species, with as many as 77% (20 species) boreo-lusitanian. this is in contrast to the records from the late weichselian, when the arctic–boreal elements dominated, with 54% in the vendsyssel region and 55% in the skagen well, the latter with only a few species and representing a deeper-water environment. the preboreal–boreal north sea faunas contain eulittoral as well as shallow-water species. mytilus edulis, littorina littorea and cerastoderma edule characterise the littoral zone and macoma balthica the shallow-water zone. using the characteristic species from the c.g.j. petersen community concept, the oldest recorded faunal communities from the north sea might be the mytilus epifauna community with littorina littorea, and the macoma infauna community with cerastoderma edule. also the abra community on mixed bottoms with phaxas pellucidus, corbula gibba and mya truncata might be reflected in the recorded species. the skagen region, region 7 appendix 6 and fig. 103 as discussed earlier, the environment of the earliest holocene, the preboreal and boreal, can be referred to the maldane-ophiura sarsi community. this deep6. vendsyssel mytilus/zirphaea z. crispata littoral arctic macoma m. calcarea shallow ophiocten zone portlandia arctica deeper 7. skagen littoral shallow arca-astarte bathyarca glacialis deeper region community species depth table 5. late weichselian communities and/or characteristic molluscan species geus bulletin no 3.pmd 09-07-2004, 09:10159 160 water community observed in the skagen region compared to the shallow-water communities recorded from the north sea coastal region once more demonstrates the unique position of the deeper-water communities observed in the skagen well material compared to the other regions in denmark through the late quaternary. the recordedpreboreal–boreal communities and/or characteristic molluscan species are given in table 6. the atlantic stage 8000–5000 14c years b.p. the bælt sea, region 1 appendix 6 and fig. 103 out of the 47 species known from the bælt sea region, 31 have been recorded from the atlantic (fig. 103). the littoral zone with mytilus edulis, littorina littorea, littorina saxatilis, cerastoderma edule, and macoma balthica from the shallow-water zone, is recorded; furthermore, the abra alba community together with corbula gibba. these faunal elements reveal the mytilus epifauna community with ostrea edulis, which no longer occurs in this area. in the infralittoral zone is the macoma infauna community where such species as paphia aurea, tapes decussatus, and venerupis pullastra have been found, and finally the abra community with corbula gibba. also the epifauna on the vegetation is reflected in rissoa albella, r. membranacea, and r. inconspicua and other gastropods. also the bivalve parvicardium exiguum is associated with the vegetation. the communities mentioned are still to be found in the bælt sea region, whereas paphia, tapes, venerupis and ostrea are no longer found in this region. the baltic, region 2 appendix 6 and fig. 103 fifteen mollusc species out of the 19 recorded from the baltic region during the holocene can be referred to the atlantic (fig. 103). theoccurrencesofmytilusedulisandlittorina littorea are referred to the mytilus epifauna community in the littoral zone, while species, as macoma balthica, cerastoderma edule, and scrobicularia plana represent the macoma infauna community in shallow water. the occurrences of both littorina littorea and scrobicularia plana are characteristic for the atlantic in the baltic and are now absent. spärck (1950) points to the wider extent of scrobicularia plana in the stone age as a consequence of warmer water in those days; however, in the present work the higher salinity is preferred as an explanation, as mentioned earlier. this is supported by the occurrences of gastropods like bittium reticulatum, rissoa albella, and aporrhais pespelicani, species recorded from other regions today with higher salinity. it should be noticed that scrobicularia plana and littorina species have been demonstrated as far north in the baltic as estonia (kessel & raukas 1979, fig. 9), although only with a low percentage but persisting into the subboreal. the unexpected find of ostrea edulis from estonia has later been re-evaluated as transported there by some seamen and thrown then overboard (in a letter from prof. a. raukas, may 1995). according to nordmann (1903b, 1906), madsen (1944), and spärck (1942, fig. 21), the southernmost finds of subfossil oysters are the bælt sea and øresund off landskrona. the kattegat, region 3 appendix 6 and fig. 103 from the kattegat region, only half of the recorded species have been dated so as to give a first appearance date. this amounts to 23 species from the atlantic (fig. 103). all of the dated species come from geologically mapped areas and not from the kattegat proper. therefore the observed species all come from shallow-water environments, excluding the deeper-water environment recorded from foraminifera (christiansen et al. 1993; seidenkrantz & knudsen 1993). 5. north sea mytilus m. edulis littoral macoma m. balthica shallow abra a. alba deeper 7. skagen littoral shallow maldana/ophiura > 150 m region community species depth table 6. preboreal–boreal communities and/or characteristic molluscan species geus bulletin no 3.pmd 09-07-2004, 09:10160 161 from the listed species dated to the atlantic the mytilus epifauna community with littorina littorea and the macoma infauna community with cerastoderma edule and tapes decussatus can be pointed out. furthermore, the abra community with corbula gibba, which is common in present-day inner danish waters (thorson 1950), is present. the limfjord, region 4 appendix 6 and fig. 103 this region has the highest number of recorded mollusc species from the holocene, viz. 147 species, and 77 have been dated to the atlantic (fig. 103). petersen (1918, pp. 22–36) described the communities in the limfjord region covering the macoma balthica, the venus and the abra communities. also an area with mya truncata is mentioned, forming a transition zone between the macoma and the abra communities. in patches the epifauna elements such as mytilus edulis and modiola modiolus are found. from the zostera vegetation, the rissoa and bittium species are mentioned. all of the characterising species from these communities have been recorded from the atlantic. paphia aurea, tapes decussatus, venerupis rhomboides and venerupis pullastra were represented during the atlantic, where as only venerupis pullastra is present in the limfjord today. ostrea edulis was well established during the atlantic, as seen from the species composition of the ‘køkkenmødding’ (kitchen midden) (petersen 1986a, figs 3, 4). the above-mentioned species and cerastoderma edule occurred in the infralittoral zone in large quantities most probably in the tidal zone which was the best collecting grounds for the stone age people. in a multi-lobed body of water such as the limfjord, many habitats have existed during the atlantic. however, also the development through time has been considered, as seen in the case of the marine stages in tastum sø – once the southernmost part of skive fjord (rasmussen & petersen 1980). in the northern part of the former limfjord during the atlantic, the deeper-water fauna with abra alba and corbula gibba can be demonstrated at the vust locality (petersen 1981, p. 502). the recorded atlantic communities have very much in common with the recent communities. the north sea, region 5 appendix 6 and fig. 103 twenty-seven species immigrated during the atlantic in the north sea coastal region (fig. 103), and chamelea striatula and spisula subtruncata characterise the venus community and are very common in the recent north sea region. also paphia aurea and tapes decussatus make their appearance in the north sea region during the atlantic. according to hessland (1943), tapes decussatus should immigrate to the west coast of sweden already in the boreal, while paphia aurea, venerupis rhomboides and venerupis pullastra followed in the atlantic. this is a close parallel to the recorded immigration to the limfjord region, although here following the transgression and not superjacent to older marine deposits as in the north sea. the vendsyssel region, region 6 appendix 6 and fig. 103 the dates from the vendsyssel region during the holocene are made on only a few species. however, the faunal assemblages sensu nordmann (jessen 1905) can be commented upon in the light of immigration dates observed in the neighbouring limfjord region. nordmann (jessen 1905, p. 145) operates with five assemblages (from a to e) with the following headings: a) beach deposits b) oyster banks c) deposits in coves and sounds d) deposits in fjords and sounds with muddy bottoms and no current e) lagoonal deposits a) beach deposits the first type – the beach deposits – cannot be considered in any relation to the community concept sensu c.g.j. petersen, since the dominating part of the shell material has been redeposited. however, as a geological unit, it points to a former sea level stand, albeit difficult to date, because of the allochthonous character of these deposits. among the 90 species listed from this region (southgeus bulletin no 3.pmd 09-07-2004, 09:10161 162 ern part of vendsyssel), nordmann (jessen 1905, table a) points to spisula subtruncata and fabulina fabula as being conspicuous, but other species may dominate at some localities, as seen from the table. both species pointed out by nordmann are recorded from the limfjord during the atlantic. b) oyster banks as pointed out by petersen (1918, p. 52), the so-called oyster banks in the recent limfjord have 1 or 2 specimens per m2. however, the places recorded by nordmann are located on former narrow channels where the oysters occurred in large quantities together with chlamys varia, hiatella arctica, retusa truncatula, mysella bidentata, parvicardium exiguum, paphia aurea, venerupis pullastra, bittium reticulatum, rissoa inconspicua, rissoa parva, buccinum undatum, nucula nitidosa, triphora adversa, cerastoderma edule, and hydrobia ulvae, all of which are recorded from the limfjord during the atlantic. only two species mentioned by nordmann (jessen 1905, p. 147) as being characteristic from some of the oyster banks, caecum glabrum and acmaea virginea, have their earliest record from the limfjord in the subboreal. therefore the oyster banks sensu nordmann seem to be well established in the vendsyssel region already during the atlantic, considering the dates obtained from the limfjord region. the oyster banks from the atlantic appear to be characteristic features with their high diversity of species and huge quantities of ostrea edulis not met with in present-day danish waters. this could be seen as a parallel to the fluctuation in the population of oysters observed during the last hundred years in danish waters, but should rather be connected with changes in the tidal currents which changed to a minimum during the following stage – the subboreal (petersen 1993), and put an end to the large oyster banks. c) deposits in coves and sounds from these deposits nordmann points to species such as spisula subtruncata, modiolus modiolus, thracia phaseolina, and corbula gibba as being characteristic of coves and sounds. they have all been recorded from the limfjord during the atlantic, and they represent species known from the deeper-water deposits both as epifaunal elements (modiolus modiolus) and infaunal elements as found in the c.g.j. petersen communities, the modiola and abra communities respectively. this is further demonstrated by the following species mentioned by nordmann (jessen 1905, p. 148): cerastoderma edule, parvicardium scabrum, nucula nitidosa, hiatella arctica, chamelea striatula, timoclea ovata, venerupis pullastra, fabulina fabula, tellimya ferruginosa, lunatia alderi, and retusa truncatulus. also these species have been dated back to the atlantic in the limfjord region. ostrea edulis occurs, but as stray finds among the infauna elements dominating in the above-mentioned assemblage that includes abra alba, which occurs in most of the samples, although not frequently (jessen 1905, table c). d) deposits in fjords and sounds with muddy bottoms and no current from such deposits nordmann mentioned the finds of zostera, which was a well established vegetational element in the recent limfjord, according to petersen & jensen (1911, map 1). the dominating species in this assemblage, which resembles the present-day fauna in such environments, are hydrobia ulvae, littorina littorea, littorina obtusata, rissoa membranacea, cerastoderma edule, mytilus edulis, scrobicularia plana, paphia aurea, bittium reticulatum, hinia reticulata, onoba semicostata, parvicardium exiguum, macoma balthica, and ostrea edulis, the last two species only with a few specimens. all the above-mentioned species occurred in the limfjord region during the atlantic. the littorina, rissoa, and parvicardium species might often be found on the zostera vegetation. among the dominating species also mentioned by nordmann, some have not been dated back to the atlantic (in the limfjord region) but occur in the subboreal, viz. littorina tenebrosa, akera bullata, and retusa obtusa. however, already the species recorded from the atlantic point to the so-called echinocyamus community (spärck & lieberkind 1921), although the echinoids have not been recorded by nordmann (jessen 1905). e) lagoonal deposits these deposits represent two assemblages, according to nordmann (jessen 1905, p. 150), viz. an older more open-water environment with species such as mactra stultorum, tellimya ferruginosa, chamelea striatula, fabulina fabula, ensis ensis, lunatia catena, lunatia alderi, and aporrhais pespelicani, which are mixed with faunal elements from the lagoon itself, such as hydrobia ulvae, scrobicularia plana, and mytilus edulis. in connectionwithazostera vegetation, rissoa membranacea and lacuna vincta may occur in huge quantities. geus bulletin no 3.pmd 09-07-2004, 09:10162 163 such a deposit cannot be compared to any of the petersen communities, although they play an important role in the geological setting, as was the case also with the beach deposits. in the northern and eastern part of vendsyssel, further comments will be added to the shallow-water and beach deposits with the finds of the dosinia and mya arenaria species. they have been dated to the subboreal and subatlantic respectively and are therefore commented upon later. the skagen region, region 7 appendix 6 and fig. 103 nearly all of the eight recovered species from the atlantic (fig. 103) show a deep-water fauna, which on the basis of the dominating role of the echinoids is tentatively referred to the amphiura community known from the present-day skagerrak. the final large eustatic rise took place during the late boreal – early atlantic, and the difference in isostatic rebound from 8000 b.p. between the skagen and limfjord regions is around 31 m, with the highest amount in the north (skagen). it appears that the water depth in the skagen region must have been up to 100 m during the atlantic (petersen 1981, 1991b). therefore, the occurrence of a single spisula subtruncata shell must be taken as far outside its habitat, considering that the modern depth range of this species is 0–36 m (petersen 1986c, table 2). the recorded atlantic communities and/or characteristic molluscan species are given in table 7. the subboreal stage 5000–2500 14c years b.p. the bælt sea, region 1 appendix 6 and fig. 103 there are no dated mollusc finds from the subboreal in the bælt sea region. as stated for the atlantic in this region, the bottom communities known from the present day were already established, but they included some species such as tapes and ostrea which are no longer extant in this area. however, as paphia aurea and ostrea edulis still occurred in the iron age sites – from the subatlantic – it is most probable that these species persisted there, while tapes decussatus and venerupis pullastra expired during the subboreal in the bælt sea region (petersen 1985c, fig. 5). the baltic, region 2 appendix 6 and fig. 103 there is no dated record of molluscs from the subboreal in the danish part of the baltic. therefore the change in the atlantic littorina fauna, into the lymnaea sea fauna, which occurred during the subboreal around 4000 b.p. (fredén 1980, p. 70), must be taken from observations outside denmark. the mollusc fauna from estonia shows that littorina littorea, rissoa membranacea, and scrobicularia plana persisted there until about the end of the subboreal (kessel & raukas 1979, fig. 9). the implications of this should be that these species must have been present in the danish area throughout the subboreal. in the central part of the baltic, around gotland, lymnaea peregra f. baltica re-immigrates after the 1. bælt sea mytilus m. edulis/ostrea edulis littoral macoma m. balthica/tapes spp. shallow abra a. alba deeper 2. baltic mytilus m. edulis/littorina littorea littoral macoma m. balthica shallow deeper 3. kattegat mytilus m. edulis littoral macoma m. balthica/tapes spp. shallow abra a. alba deeper 4. limfjord mytilus/modiola m. edulis/tapes spp. littoral macoma m. balthica shallow abra/venus a. alba deeper 5. north sea littoral shallow venus chamelea striatula deeper 6. vendsyssel littoral modiola m. modiolus shallow abra a. alba deeper 7. skagen littoral shallow amphiura parvicardium minimum ~ 100 m region community species depth table 7. atlantic communities and/or characteristic molluscan species geus bulletin no 3.pmd 09-07-2004, 09:10163 164 maximum of the littorina transgression (munthe 1940, p. 124). this gastropod was also present in the early, more brackish part of the littorina sea deposits in fakse bugt. in estonia (kessel & raukas 1979, fig. 9), the reappearance of lymnaea took place around 4000 b.p., implying that the salt-demanding species (littorina, rissoa, and scrobicularia) occurred together with the brackish lymnaea species throughout the later part of the subboreal! the kattegat, region 3 appendix 6 and fig. 103 the mollusc faunas recorded from the kattegat region represent only part of the total faunal complex within this large region, and have been dated only on djursland. however, this demonstrates the expiring tidal amplitude in the early part of the subboreal. the only dated immigrants to the fauna from the central part of djursland during the subboreal are onoba semicostata and littorina tenebrosa. both extended into the baltic today and tolerate brackish water. in this way they are typical for the environmental changes recorded in the marine faunas from djursland. the fauna during the atlantic was characterised by ostrea edulis, tapes decussatus, macoma balthica, and corbula gibba. bittium reticulatum was present in large quantities, but disappeared in the subboreal. also the decline in numbers of hydrobia ulvae and its replacement in equal numbers by hydrobia ventrosa speak in favour of a more brackish-water influence. the implications of the study of mollusc species on a quantitative basis in connection with 14c dates and pollen analyses confirm that the fauna during the tapes sea period was more prolific than nowadays. however, it also demonstrates as a new point of view that this applies only for the atlantic. in petersen (1993, p. 368) it is argued that the change in sedimentation rate from the atlantic to the subboreal, which has been calculated for the korup sø area on djursland, points to a lowering of the tidal range in danish waters since the atlantic. this is explained in that way that sedimentation will stay low as far as the tidal current reaches and allows halophilous species to live far up in the fjords according to the observations on the faunal record. furthermore, an older record from the mapping of the area of flaser bedding seen as a tidal bedding supports such an explanation. it was tempting to see the change from the littorina sea to the lymnaea sea in the baltic on the background of such a lowering of the tidal impact in the inner danish waters. however, as shown in the preceding section on the baltic, the change occurred around 4000 b.p. recalling the statement by c.g.j. petersen that the deposition of the tapes layers has happened in a period when the danish waters from a hydrographical point of view have been more like the north sea or the open sea than now, it is clear that a tidal impact could make the difference and explain the large oyster banks far into the roskilde fjord in north-eastern sjælland and other former fjord regions facing the kattegat region. the well-dated ertebølle coastal sites (‘køkkenmøddinger’ – kitchen middens) from all over denmark also present a large amount of ostrea edulis from the atlantic and demonstrate that the molluscan diet later in the subboreal was based on the cardium species (andersen 1991, 1995). this situation has lasted into the iron age, as seen in the shell middens from the bælt sea area (petersen 1985c, fig. 5). however, this change mostly affected the fjord complex. consequently the kattegat region still has the communities listed for the atlantic. the limfjord, region 4 appendix 6 and fig. 103 the 36 species which immigrated into the limfjord during the subboreal (appendix 6) can be considered according to their way of life, presented from the list below. age: subboreal climatic regions: asb. class gastropoda subclass prosobranchia order archaeogastropoda margarites helicinus (phipps 1774) subclass opisthobranchia order anaspidea retusa obtusa (montagu 1803) total for climatic regions asb. : 2 (5.6%) climatic regions: asbl class gastropoda geus bulletin no 3.pmd 09-07-2004, 09:10164 165 subclass prosobranchia order archaeogastropoda acmaea tessulata (müller 1776) order neotaenioglossa lacuna pallidula (da costa 1778) total for climatic regions asbl: 2 (5.6%) climatic regions: .sbl class gastropoda subclass prosobranchia order archaeogastropoda acmaea virginea (müller 1776) class bivalvia subclass pteriomorpha order pterioida delectopecten vitreus (gmelin 1791) total for climatic regions .sbl: 2 (5.6%) climatic regions: ..bl class gastropoda subclass prosobranchia order archaeogastropoda helcion pellucidum (linnaeus 1758) gibbula tumida (montagu 1803) skenea basistriata (jeffreys 1877) order neotaenioglossa littorina tenebrosa (montagu 1803) lacuna parva (montagu 1803) cingula semistriata (montagu 1808) rissoa violacea desmarest 1814 caecum glabrum (montagu 1803) aporrhais pespelicani (linnaeus 1758) order heterogastropoda epitonium turtonis (turton 1819) vitreolina philippii (rayneval & ponzi 1854) order neogastropoda hinia incrassata (ström 1768) raphitoma purpurea (montagu 1803) raphitoma linearis (montagu 1803) subclass heterobranchia order heterostropha chrysallida decussata (montagu 1803) chrysallida indistincta (montagu 1808) ebala nitidissima (montagu 1803) eulimella laevis (brown 1827) ondina divisa (j. adams 1797) ondina diaphana (jeffreys 1848) subclass opisthobranchia order anaspidea akera bullata müller 1776 class bivalvia subclass palaeotaxodonta order nuculoida nucula nucleus (linnaeus 1767) subclass pteriomorpha order pterioida palliolum striatum (müller 1776) palliolum tigerinum (müller 1776) subclass heterodonta order veneroida mactra stultorum (linnaeus 1758) lutraria lutraria (linnaeus 1758) ensis ensis (linnaeus 1758) order myoida pholas dactylus linnaeus 1758 total for climatic regions ..bl: 28 (77.8%) climatic regions: ...l class gastropoda subclass prosobranchia order neotaenioglossa alvania lactea (michaud 1830) onoba proxima (forbes & hanley 1850) total for climatic regions ...l: 2 (5.6%) total for age subboreal: 36 (14.6%) the archaeogastropoda are all (six) epifauna on seaweeds or on hard substrates. the neotaenioglossa have seven epifaunal elements mostly on seaweeds and two infauna species, of which aporrhais pespelicani is a shallow infauna animal. the heterogastropoda with epitonium turtonis and vitreolina philippii are associated with other animals, the former feeding on anemone or preying on other species and the latter being an intermittent parasite of echinoderms (fretter & graham 1982, p. 387). the heterostropha with six species are predators or external parasites. the anaspida with two species, akera bullata and retusa obtusa, are epifauna and infauna species respectively, the former on zostera in shallow water and the latter in mud or fine sand connected with the macoma community. the only nuculoida found, nucula nucleus, belongs to the shallow infauna. the pterioida with three species are referred to the epifauna, since the delectopecten vitreus is found attached with its byssus on hard substrates. the veneroida have three species which are all referred to the infauna. lutraria lutraria and ensis ensis are deep-borrowing. geus bulletin no 3.pmd 09-07-2004, 09:10165 166 the myoida with pholas dactylus bores in different substrates. when taking the above-mentioned groups of species associated with other animals, carnivores, predators, and boring species as a whole, we have three categories: the epifauna with 47.2%, the infauna with 19.4%, and other elements with 33.4% of the species immigrated during the subboreal. when the same procedure is followed for the 77 species which have been dated to the atlantic in the limfjord region, we find that the percentages for the epifauna, the infauna and other elements are 31%, 46.8% and 22% respectively, which shows that the epifauna element becomes the dominating one in the subboreal among the newcomers. this might tentatively be connected with a denser vegetation in the subboreal of sea-weed. the north sea, region 5 appendix 6 and fig. 103 nineteen species make their first appearance in the danish north sea during the subboreal. when considering their way of living and their grouping into epifauna, infauna and other elements, it appears that the groups are of equal size, i.e. five, six and eight species respectively. however, the number is too low to be used for any comparison with other regions. in the north sea region the dates of first appearance go back to the preboreal–boreal stage, showing that the initial stages were dominated by the infauna species; the preboreal–boreal: seven epifauna, 16 infauna, and four other elements; the atlantic: six epifauna, 17 infauna, and four other elements. the development of the bottom communities in the north sea region seems in this way to corroborate the changes observed in the limfjord region from the atlantic to the subboreal. these changes are in facies rather than climatic. a slightly more temperate fauna was met with during the atlantic, as mentioned earlier, and it has consequently no bearing on the observed changes. but the expiring tidal influence in the danish waters taking place in the early subboreal might have been of some importance for the environmental changes reflected through the bottom communities. the vendsyssel region, region 6 appendix 6 and fig. 103 the vendsyssel region does not give much information on the immigration of species during the holocene. however, the dosinia beds were described from this area and have been dated quite recently in the type area around strandby north of frederikshavn (nordmann 1904; petersen 1991b). the oldest date for the dosinia exoleta, which is the characteristic species for the dosinia beds, is 4240 ± 85 b.p. in 14c years (k5318). this earliest dated occurrence of dosinia exoleta corresponds to a hydrographical change in the kattegat region described by nordberg & bergsten (1988) and nordberg (1989). the demonstrated lowering of the tidal influence in the inner danish waters took place also in the early part of the subboreal. petersen (1976) pointed out that seven mollusc species hitherto known only from the dosinia beds also occur in the raised marine deposits from the western part of the limfjord, i.e. lucinoma borealis, hinia incrassata, venerupis rhomboides, abra prismatica, lutraria lutraria, pholas dactylus, and helcion pellucidum. furthermore, not only ten species from the dosinia beds are also in the deposits from the limfjord but five of them occurred already during the atlantic: gari fervensis, turritella communis, lucinoma borealis, abra prismatica, and venerupis rhomboides. these species, representing an infauna assemblage very much like the dosinia species, were also characteristic of the early holocene dominating infauna mollusc assemblage. mörner (1969, pp. 384–386, and table 1) points out that some species in the dosinia fauna occur in older layers along the swedish west coast, referring to the works by hessland (1943) and antevs (1917). however, this is not the case with the characterising dosinia species, in as much as dosinia exoleta has not been demonstrated in the studies by hessland and antevs and dosinia lincta occurs only in the younger deposits referred to the subboreal. among the 15 species listed, eight have been recorded from the limfjord, out of which epitonium turtonis, oenopota turricola, acteon tornatilis, and cylichna cylindracea have their first appearance in the atlantic, and lutraria lutraria, pholas dactylus, and alvania lactea appeared in the subboreal, whereas the occurrence of modiolus adriaticus in the limfjord has not been dated. in this way, we are left with only six species which have not been found in other regeus bulletin no 3.pmd 09-07-2004, 09:10166 167 gions outside vendsyssel older than the subboreal: pecten maximus, dosinia exoleta, dosinia lincta, gari depressa, alvania cimicoides, and trivia monacha. only the characterising species dosinia exoleta has been dated as mentioned above, and recently donax vittatus from vr. holmen in the northern part of vendsyssel, west of strandby, with the oldest date of this bivalve hitherto obtained in the danish deposits, 4240 ± 75 14c age b.p. (aar-1481). this date shows that donax characterising the highenergy shore deposits occurred in danish waters since the subboreal. the further revision of the dosinia fauna shows that only very few species are limited to the vendsyssel region both in time and space. therefore, it cannot be sustained for the danish material as mentioned by mörner (1969, p. 384) that: “the dosinia layers contains a great number of new boreo-lusitanic immigrants” of the molluscs entirely belonging to the dosinia layers according to nordmann (1904), only three species, out of the 26 species mentioned, do not occur in danish waters today, according to jensen & knudsen (1995), viz. trivia monacha, gari depressa, and alvania cimicoides. only trivia monacha seems to be purely lusitanian, since a recent distribution to the north sea is questioned by fretter & graham (1981, p. 329), and there is no record from scandinavia. the other two species have a boreo-lusitanian distribution. the skagen region, region 7 appendix 6 and fig. 103 the 23 species, out of which only vitreolina collensi is purely lusitanian, can be evaluated according to their way of life. vitreolina collensi belongs together with aclis minor and melanella alba to the eulimacea, which are probably associated with echinoderms (fretter & graham 1982, p. 397). the eulimacea, together with the epitoniacea, usually prey on anthozoans. the heterostropha, including the family pyramidellidae which lives ectoparasitically on other marine organisms (jensen & knudsen 1995), are here represented by eulimella scillae. finally within the veneroidae, mysella bidentata and tellimya ferruginosa are commensals on echinoids, but can also be found free-living. for the rest of the 23 species found during the subboreal in the skagen well, it applies that 16 species belong to the infauna, including onoba vitrea which tends “to live in muddy places, often so muddy that one wonders how the animals keep the mantle cavity clear” (fretter & graham 1978b, p. 170). it appears from the above-mentioned dates based on type of bottom-dwelling animals that the fauna belongs to the deep-level sea bottoms which goes well together with the turritella-venus communities. the recorded subboreal communities and/or characteristic molluscan species are given in table 8. the subatlantic stage 2500– 14c years b.p. the configuration of land and sea in the danish realm was very close to that of today. the isostatic movements during this time span have been so small that they did not affect the general outline (petersen 1991b). however, the coastal development, in the formation of simplified coastline and spits especially in the west towards the north sea and in the north facing the skagerrak and the kattegat respectively, still affected the contour of the land. 1. bælt sea mytilus m. edulis littoral macoma m. balthica/ostrea/paphia shallow abra a. alba deeper 2. baltic mytilus m. edulis littoral macoma m. balthica shallow deeper 3. kattegat mytilus m. edulis littoral macoma m. balthica shallow abra a. alba deeper 4. limfjord mytilus/modiola m. edulis littoral macoma m. balthica shallow venus/abra a. alba deeper 5. north sea littoral shallow venus chamelea striatula deeper 6. vendsyssel donax d. vittatus littoral dosinia d. exoleta shallow deeper 7. skagen littoral shallow venus/turritella t. communis deeper table 8. subboreal communities and/or characteristic molluscan species region community species depth geus bulletin no 3.pmd 09-07-2004, 09:10167 168 for the main part of the danish waters, the recent marine bottom communities were established, some of them already since the atlantic, although the few characterising tapes and ostrea species are no longer extant in great numbers or have totally vanished from the danish seas. therefore, the actual map of the petersen (1914, 1918) bottom communities as seen today will be commented upon in relation to the few, but important changes observed during the subatlantic, region by region. the bælt sea, region 1 appendix 6 and fig. 103 the bottom communities mapped from the bælt sea region comprise the macoma balthica community in the shallow-water area and the abra alba community in deeper water (petersen 1918), the latter community with tridonta borealis and macoma calcarea. the former has been recorded from the subfossil finds but not dated, while the latter has a dated occurrence back in the atlantic and is considered part of the abra alba community as a deep infauna element. considering the present distribution of the astarte species, it is most probable that tridonta borealis invaded the bælt sea and the baltic already in the early holocene along with the transgression in the early atlantic. the gregarious occurrences of ostrea edulis recorded from the atlantic in the bælt sea region vanished in the iron age (petersen 1985c). this species is no longer found in the bælt sea region, nor is paphia aurea, which also occurred at the iron age sites too (petersen 1985c, fig. 5). the steady occurrence of ostrea edulis since the atlantic, although in reduced numbers, might have led to an experiment in cultivating oysters south of lolland in the fehmern bælt (winther 1876, p. 114), although an unsuccessful one. the distribution of oysters within the danish waters seems to have changed very much right up to the present day, with many records from the 19th century of oyster banks from places where no records are found today (kröyer 1837; seaman & ruth 1997). the baltic, region 2 appendix 6 and fig. 103 the macoma balthica community covers the whole area of the baltic, implying that in this area the otherwise shallow-water bivalve extends into greater depths – more than 50 m (petersen 1918). also in this area, tridonta borealis and tridonta elliptica have been found in great quantities east of bornholm but not dated. however, according to johansen (1916, fig. 5), tridonta borealis and tridonta elliptica are recorded only from areas with a salinity of more than 10‰, but zenkevitch (1963, p. 338, fig. 167) points to many finds further to the north in the baltic, where the salinity is lower. as stated earlier, the change from the littorina sea stage to the lymnaea sea stage took place during the subboreal. the present situation with a mya sea stage – a term established by munthe (1894) – took place at a very late date. munthe (1894, p. 14) said: “since mya arenaria is an easily identified and characteristic species in the present baltic it seems suitable to call the present time the ‘mya-time’ or ‘mya arenaria-time’ in opposition to ‘littorina-time’ etc.” the kattegat, region 3 appendix 6 and fig. 103 among the subfossil species both, dated and undated, no records of chlamys striatula and turritella communis are found. these characterising species for the venus and deeper venus communities respectively have a wide extension on the map by petersen (1918) in the kattegat region. also the deeper-water epifaunal elements – characterised by modiolus modiolus – are missing in our subfossil record. only the macoma balthica and abra alba communities are recognised in the subfossil material. however, the development in the skagen well sequence to the north in the kattegat region of mollusc species reveals the venus-turritella communities and can be taken as part of the development in the central kattegat region not sampled at the time of this study on molluscs. in the present day northern kattegat, stray specimens of ostrea edulis have been recorded (jensen & knudsen 1995, p. 40). otherwise among the more spectacular tapes sea species, tapes decussatus, dated from the atlantic, and paphia aurea, not dated but occurgeus bulletin no 3.pmd 09-07-2004, 09:10168 169 ring in the subfossil fauna, have disappeared from the danish waters. the limfjord, region 4 appendix 6 and fig. 103 this region has been studied in more detail, regarding the recent fauna, than the other regions, and references can be made also to jensen (1919). the abra community is here divided into three associations, i.e. nucula-corbula, abra-solen, and abrasolen-mya associations. in more shallow water the abra community is replaced by the macoma balthica community. all of these communities are recorded by their mollusc species in the subfossil fauna, here including the mya arenaria in the subfossil assemblage. however, the only dated subfossil immigrating species from thesubatlantic isdonaxvittatus,whichappearedaround 2000 years b.p. in the northernmost part of the limfjord region in the beach ridges, around 1000 years before the closing of the western and northern entrance to the limfjord. the closing of the entrances from the north sea and skagerrak changed the limfjord region into a freshwater basin between a.d. 1200 and 1825, however, with periods of saltwater influence (kristensen et al. 1995). a comparison between the subfossil fauna before the freshwater stage and the recent one after the north sea broke through in the western part of the limfjord in 1825 shows that the subfossil fauna had only a slightly higher affinity to more temperate water than the recent one. also in this place the tapes, venerupis and paphia species make the difference, in the way that only venerupis pullastra finds its way back to the region after 1825. paphia aurea, however, has a dated occurrence from the same deposits as donax vittatus to 1910 ± 100 14c years b.p. (petersen 1976). it is seen that paphia aurea in this region, as in the bælt sea, has a record up into the iron age before it became extinct in the danish waters. ostrea edulis repopulated the limfjord region after 1825 and reached a wide extension in this region already in the second half of the 19th century (collin 1871). however, the population has suffered from strong fluctuations not only in the limfjord but also in other danish waters, as shown on the map by kröyer (1837). spärck in several papers on the biology of oysters (ostrea edulis), published in reports of the danish biological station, also discussed the fluctuations in the nw european population of oysters (spärck 1950, pp. 43–45). spärck reached the conclusion that the summer temperature of the water was crucial, both being too low and too high, which affected the oyster in its reproduction and in food supply respectively. furthermore, severe winters might affect the population, although less than the summer temperatures. however, these changes did not mean a total disappearance of the oyster, but only a reduction to such a level that the industrial exploitation had to stop. when taking into account the many studies on the population of oysters, one could use the results in a general conclusion on the variations found in the whole population of molluscan species, especially for the group having their northern limit within the boreal region: even small variations in the climate may influence the size of the population. also the environmental changes as shown within the danish area during the holocene, such as the lowering of the tidal amplitude in the early part of the subboreal, had a severe influence on the populations in the inner danish waters. here again, oysters can be taken as an example by the termination of the huge oyster banks known from the atlantic. spärck (1950, p. 44) draws attention to the oyster banks in holland and the british isles, where the density of the population is far greater because of the tidal movements. however, not only the hydrographical changes through time, in the tide, but also the coastal evolution, such as the formation of simplified coastline and spits, play an important role in the distribution and new finds of molluscan species. the north sea, region 5 appendix 6 and fig. 103 only eight species have been recorded as immigrants during the subatlantic. however, two of them, donax vittatus and dosinia lincta, deserve special attention. only parvicardium ovale and dosinia lincta have their first dated appearance. the other species, except donax vittatus with occurrence in the subboreal, have been recorded from the atlantic at various places listed in appendix 6. in referring to the c.g.j. petersen bottom community map covering also the north sea, the macoma geus bulletin no 3.pmd 09-07-2004, 09:10169 170 balthica and the venus communities are found in the danish north sea coastal region, the former in bays and off the southern part of the west coast (petersen 1914), the latter around the westernmost part of the limfjord and the jydske rev wnw of the bovbjerg coastal cliff (petersen 1994a). a landscape like the limfjord of today was found 75 km towards wnw in the area of the jydske rev. following the transgression in the early part of the holocene, the glacial landscape in an area of the present jydske rev was eroded and the high-energy coast approached the appearance of the present one. in the northern part forming an erosion coast and in the southern part at blåvands huk an aggradation coast, both characterised by the presence of donax vittatus. in the southern part the aggradation started around 800 bc some 2000 m east of the present coastline (petersen 1994a, p. 24) as seen from the dating of donax vittatus to 2620 ± 75 b.p. 14c years (aar-1480) off the inland cliff at grærup (fig. 1). at bovbjerg, the strata with donax vittatus in the agger spit are dated to 410 ± 65 b.p. (petersen 1985a). the formation of the spits closing the former bays on the jylland west coast is a consequence of the formation of a simplified coast. further to the north, donax vittatus from kovad bro in the northernmost part of the limfjord, 6 km inland, gave a date of 1910 ± 100 b.p. (petersen 1976), showing that the beach progressed 6 km during approximately 2000 years (petersen & andreasen 1989, fig. 1). it is tempting to introduce the idea that the enormous change in the land–sea configuration in the eastern part of the north sea affected the tidal currents in the inner danish waters. this could possibly have occurred when most of the jydske rev formation was eroded to such a level that the tidal current from the south was no longer braked and consequently the present-day interference with the tidal current came into existence. it is the interference between the two tidal currents in the skagerrak today that makes the tidal amplitude small in the inner danish waters (nielsen 1939; kuenen 1950). dosinia lincta has been dated (870 ± 110 b.p. 14c years) in the jydske revsand formation in the vibrocore 562001 around 75 km off the coast of jylland at a depth of 32 m (petersen 1994a, p. 18, fig. 3). the assemblage from these strata comprises spisula subtruncata, phaxas pellucidus, fabulina fabula, chamelea striatula, dosinia exoleta, corbula gibba, cochlodesma praetenue and thracia phaseolina, most of them characterising the jydske revsand formation. the vendsyssel region, region 6 appendix 6 and fig. 103 the coastal development in the eastern part of the region facing the kattegat takes place in the form of migrating bars (schou 1949, fig. 17b), the so-called rimmer and doppe system sensu jessen (1905). the venus and the macoma balthica communities are found in the coastal zone, the former dominating in the northern part, whereas the latter forms a small area between shore and the venus community to the south towards the entrance to the limfjord at hals (petersen 1918). in the north at strandby, locus typicus of the dosinia beds, the layers with dosinia exoleta are superposed by a layer characterised by spisula subtruncata. these beds with spisula in great quantities were dated to 2640 ± 75 (petersen 1991b), the end of the subboreal, and at a level of 4.2 m a.s.l. this corresponds to a stage in the development of the skagen spit up to 4 km south of højen, where the beach ridges have an elevation of 5 m a.s.l. in the southern part of the vendsyssel region around hals another of the faunal elements of the dosinia beds – lutraria lutraria – has for long been regarded as extinct (petersen 1992). however, “from 1990 onwards live specimens have been collected regularly near frederikshavn and on the skagerrak-coast” (jensen & knudsen 1995, p. 43). also many shells of lutraria lutraria were found along the shore south of jerup halfway between frederikshavn and aalbæk. the immigration of mya arenaria cannot be taken as an indication of changes in climate, as this species mainly belongs to the boreal region and has been transferred by man from north america. what made the find so important has a more historical than geological bearing, namely that the dates obtained from the sampling at jerup demonstrated that the american softshell clam (mya arenaria) predated columbus’ voyage in 1492, having an age of a.d. 1245–1295 at ± 1 s.d. this led to the conclusion that the vikings were better candidates than columbus to be the first to find north america (petersen et al. 1992b). the significance of changes in facies is clearly demonstrated in the next and final section describing the subatlantic faunal development in the skagen region. geus bulletin no 3.pmd 09-07-2004, 09:10170 171 the skagen region, region 7 appendix 6 and fig. 103 the subatlantic molluscan fauna from the skagen well comprises 68 species with 75% belonging to the boreolusitanian region and only one lusitanian species, vitreolina collensi. however, the more interesting fact from the younger part of the skagen sequence is the total lack of macoma balthica. in this way it presents the finest resemblance with the recent bottom community map (petersen 1918), and shows that the macoma balthica community disappears in the northern part of the east coast of jylland. this means that during the last stage of the spit formation at the site of the present skagen animals, from the venus community dominated along shore. this is alsodocumentedby thehighamount of infauna elements, with 35 out of the 68 species recorded. furthermore, some of the 11 epifauna gregarious species usually connected with the vegetation can be excluded, since they occur only as stray finds, viz. lacuna pallidula, rissoa violacea, and bittium reticulatum, as discussed earlier. the rest of the molluscs (22 species) are carnivores, predators, external parasites, and commensals. turning these figures into percentages, the epifauna species amount to 16.2% and the infauna to 51.5%. comparing this with the limfjord region where an equal number of species have been found during the atlantic and the subboreal, it appears that the number of infauna species from the atlantic to the subboreal falls from 46.8% to 19.4% and the epifauna elements rise from 31.2% to 47.2% in the limfjord. counting the limfjord region as an inner danish water today, it is worth noticing that during the atlantic the situation was much more like the ‘open’ waters as seen in the skagen figures. considering that the tidal amplitude really was lowered in the early part of the subboreal, this would to some extent explain the observed changes in the limfjord from the atlantic to the subboreal. the recorded subatlantic communities according to the maps by petersen (1914,1918) with characteristic molluscan species are shown in table 9. table 9. subatlantic communities according to the bottom community maps* with characteristic molluscan species 1. bælt sea mytilus m. edulis littoral macoma m. balthica/÷ostrea/paphia shallow abra a. alba deeper 2. baltic mytilus m. edulis littoral macoma m. balthica/+mya arenaria shallow macoma m. balthica deeper 3. kattegat mytilus m. edulis littoral macoma m. balthica shallow abra/venus a. alba deeper 4. limfjord mytilus/modiola m. edulis littoral macoma m. balthica/÷paphia aurea shallow abra a. alba deeper 5. north sea donax d. vittatus littoral macoma m. balthica shallow venus chamelea striatula deeper 6. vendsyssel littoral macoma m. balthica shallow venus chamelea striatula deeper 7. skagen donax d. vittatus littoral spisula s. subtruncata shallow venus chamelea striatula deeper * petersen (1914, 1918). region community species depth geus bulletin no 3.pmd 09-07-2004, 09:10171 172 concluding remarks in the last section on the environmental changes within the seven regions through the late quaternary, it has been demonstrated how the skagen sequence ‘moved’ into the present-day faunal community known from this area, ‘coming’ from older deposits in many ways, according to the molluscs, different from what hitherto was known in other parts of the danish area during the late quaternary. the development is graphically shown in fig. 93 (fold-out, back cover). regarding the climatic changes, recalling figs 102 and 103, the molluscs have given a clear record as far as the main trends are concerned – the interglacial– glacial cycle. however, the climatic changes during the holocene, if they were ever more than small, were overshadowed by the facies changes affecting the danish area. from the eemian as well, it must be concluded that differences in facies made the difference between the regions, and that the well-established more temperate eemian marine fauna was connected only with the shallow-water environment. among the 140 species recorded from the eemian, 118 or 84.3% occur in the holocene subfossil material and/or recent fauna. however, it should be noticed that among the 22 species only found in the eemian, ten species or 7.1% are purely lusitanian forms, which include the no longer extant species paphia senescens. the lusitanian forms are: one gastropod – haminoea navicula; one scaphopod – dentalium vulgare; and seven bivalves – mytilaster lineatus, mytilaster solidus, lucinella divaricata, plagiocardium papillosum, gastrana fragilis, abra segmentum and gouldia minima. along with the high percentage of purely lusitanian forms – compared with and not found in the holocene – it is shown that among the 90 boreo-lusitanian species from the eemian, by far the dominating group covering 64.3% of the fauna, 87 species occur also in the holocene. the weichselian marine fauna known from the older yoldia and younger yoldia sea deposits also has a characteristic of its own, with about one third of the fauna restricted to the yoldia seas. nearly half of them are either purely arctic, such as the two bivalves portlandia arctica and macoma torelli, or with arctic– subarctic affinities, such as the three gastropods alvania cruenta, turritella erosa and cylichna occulta, and the four bivalves bathyarca glacialis, macoma loveni, pandora glacialis, and lyonsia arenosa. one third of the weichselian fauna is found both among the eemian and the holocene species (subfossil and/or recent). for the last third, the majority (11 out of 17) are also recorded only from the holocene. when the weichselian marine fauna itself is looked at, comparing the older and younger yoldia sea faunas, 23 out of the 54 species are common to both, while one third is only found in the younger yoldia sea deposits from where the subarctic–boreal species among the bivalves are: nuculana minuta, arctica islandica, and zirfaea crispata. the subarctic–boreal– lusitanian species count one polyplacophor: tonicella marmorea; one gastropod: buccinum undatum; and two bivalves: mytilus edulis and macoma balthica. this demonstrates in the best way the boreo-arctic impact around 13 000 b.p. (14c years) in the shallow-water environment characterised by zirfaea crispata and mytilus edulis – the former giving name to the deposits of that time in vendsyssel. however, the main result of this investigation was the comparison between the fossil faunas and the molluscan faunas now living before our eyes – as c.g.j. petersen expressed it in 1910 – using the c.g.j. petersen bottom community concept step by step in the seven stages from the eemian to the subatlantic within the seven regions in the danish realm. acknowledgements this study was supported by a one-year grant from the carlsberg foundation. during that year – and the following years – the geological survey of denmark and greenland (geus) has supplied me with all the facilities needed for the research. the support of both these institutions is greatly appreciated. among the many helpful colleagues at the survey, i would like to thank lasse gudmundsson for keeping order in the many samples that were analysed, and frants von platen-hallermund for making the compilations seen in the figures and appendix listing the molluscan species and other data. state geologist at that time johnny fredericia, who caused me to take up the challenge and continued to support me is thanked, as is richard bradshaw, state geologist of my new department at the survey, the department of environmental history and climate geus bulletin no 3.pmd 28-06-2004, 08:46172 173 change. j. heinemeier contributed with appendix 4 on the 14c dates on shell macrofossils from the skagen cores. kaare l. rasmussen read an early draft of the work and the referees h.g. petersen and s. funder contributed to make the work better – all are thanked. with great experience in writing for me, birgit jørgensen did the typing and commented upon the english. susanne veng christensen has made the final copy. peter john crabb revised the english in the most thorough way. to the former curator of the vertebrate collection at the geological museum, university of copenhagen, now happily at the gram museum, ella hoch, this book is dedicated. references andersen, b.g. 1965: the quaternary of norway. in: rankama, k. 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ovata acanthocardia echinata (linnaeus 1758) cardium echinatum acar nodulosa (müller 1766) aclis ascaris (turton 1819) aclis minor (brown 1827) aclis walleri jeffreys 1867 acmaea tessulata (müller 1776) acmaea virginea (müller 1776) tectura virginea acteon tornatilis (linnaeus 1758) adipicola simpsoni (marshall 1900) admete viridula (fabricius 1780) aequipecten opercularis (linnaeus 1758) pecten opercularis akera bullata müller 1776 alvania abyssicola (forbes 1850) rissoa abyssicola alvania cimicoides (forbes 1844) rissoa cimicoides alvania cruenta odhner 1915 alvania jan mayeni (friele 1886) rissoa jan mayeni alvania jeffreysi (waller 1864) alvania lactea (michaud 1830) rissoa lactea alvania punctura (montagu 1803) rissoa punctura alvania scrobiculata (möller 1842) rissoa scrobiculata amauropsis islandicus (gmelin 1791) angulus fabulus → fabulina fabula angulus tenuis (da costa 1778) tellina tenuis anomia aculeata → heteranomia squamula anomia ephippium linnaeus 1758 anomia patelliformis → pododesmus patelliformis anomia squamula → heteranomia squamula antalis agile g.o. sars 1878 antalis entalis (linnaeus 1758) dentalium entalis aporrhais pespelicani (linnaeus 1758) aporrhais serresianus (michaud 1828) arca glacialis → bathyarca glacialis arcinella plicata → saxicavella jeffreysi arcopagia crassa (pennant 1778) arctica islandica (linnaeus 1767) cyprina islandica geus bulletin no 3.pmd 28-06-2004, 08:46180 181 assiminea grayana fleming 1828 astarte banksie → tridonta montagui astarte borealis → tridonta borealis astarte compressa → tridonta elliptica astarte sulcata (da costa 1778) axinopsida orbiculata (g.o.sars 1878) axinopsis orbiculata axinopsis orbiculata → axinopsida orbiculata axinus ferruginosus → leptaxinus ferruginosus axinus flexuosus → thyasira flexuosa balcis devians → vitreolina philippii barleeia unifasciata (montagu 1803) barnea candida (linnaeus 1758) pholas candida bathyarca glacialis (gray 1824) arca glacialis bathyarca pectunculoides (scacchi 1834) bela exarata g.o.sars 1818 bela incisula → oenopota incisula bela nobilis → oenopota turricola bela trevelliana → oenopota trevelliana bela turricola → oenopota turricola bela violacea → oenopota violacea bittium reticulatum (da costa 1778) cerithium reticulatum boreotrophon clathratus (linnaeus 1767) trophon clathratus boreotrophon truncatus (ström 1768) brachystomia carozzai van aartsen 1987 brachystomia eulimoides hanley 1844 odostomia eulimoides odostomia pallida buccinum cyaneum bruguière 1792 buccinum groenlandicum buccinum groenlandicum → buccinum cyaneum buccinum undatum linnaeus 1758 cadulus jeffreysi → cadulus subfusiforme cadulus subfusiforme (m. sars 1865) cadulus jeffreysi caecum glabrum (montagu 1803) calliostoma formosa (mighels 1842) calliostoma zizyphinum (linnaeus 1758) callochiton septemvalvis (montagu 1803) capulus ungaricus (linnaeus 1758) cardium ciliatum → clinocardium ciliatum cardium echinatum → acanthocardia echinata cardium edule → cerastoderma edule cardium edule → cerastoderma glaucum cardium exiguum → parvicardium exiguum cardium fasciatum → parvicardium ovale cardium groenlandicum → serripes groenlandicus cardium minimum → parvicardium minimum cardium nodosum → parvicardium scabrum cardium norvegicum → laevicardium crassum cardium papillosum → plagiocardium papillosum cerastoderma edule (linnaeus 1758) cardium edule cerastoderma glaucum (poiret 1789) cardium edule var. balticum cerithiella metula (lovén 1846) cerithiopsis barleei jeffreys 1867 cerithiopsis tubercularis (montagu 1803) cerithium reticulatum → bittium reticulatum chamelea striatula (da costa 1778) venus gallina chemnitzia lactea → turbonilla lactea chlamys islandica (o.f.müller 1776) pecten islandicus chlamys varia (linnaeus 1758) pecten varius chrysallida eximia (jeffreys 1849) parthenina eximia chrysallida decussata (montagu 1803) chrysallida indistincta (montagu 1808) parthenia indistincta chrysallida obtusa (brown 1827) parthenia interstincta chrysallida spiralis (montagu 1803) parthenia spiralis cingula proxima → onoba proxima cingula semistriata (montagu 1808) putilla semistriata rissoa semistriata cingula striata → onoba semicostata cingula turgida (jeffreys 1870) cingula vitrea → onoba vitrea circe minima → gouldia minima claturella linearis → raphitoma linearis clausinella fasciata (da costa 1778) venus fasciata clinocardium ciliatum (fabricius 1780) cardium ciliatum clione limacina (phipps 1774) cochlodesma praetenue (pulteney 1799) colus gracilis (da costa 1778) colus jeffreysianus (fischer 1868) colus sabini (gray 1824) corbula gibba (olivi 1792) crenella decussata (montagu 1803) crepidula fornicata (linnaeus 1758) cultellus pellucidus → phaxas pellucidus cupidaria cuspidata (olivi 1792) geus bulletin no 3.pmd 28-06-2004, 08:46181 182 cyamium minutum → turtonia minuta cylichna alba (brown 1827) cylichna cylindracea (pennant 1777) cylichna occulta (mighels 1841) cylichna scalpta cylichna propinqua cylichna propinqua → cylichna occulta cylichna scalpta → cylichna occulta cypraea europaea → trivia monacha cyprina islandica → arctica islandica cytharella coarctata (forbes 1840) mangelia costata delectopecten vitreus (gmelin 1791) dentalium entalis → antalis entalis dentalium vulgare da costa 1778 devonia perrieri (malard 1904) diaphana hyalina → diaphana minuta diaphana minuta brown 1827 diaphana hyalina divaricella divaricata → lucinella divaricata donax vittatus (da costa 1778) dosinia exoleta (linnaeus 1758) dosinia lincta (montagu 1803) dosinia lupinus dosinia lupinus → dosinia lincta ebala nitidissima (montagu 1803) eulimella nitidissima emarginula fissura (linnaeus 1758) ensis arcuatus (jeffreys 1865) ensis ensis (linnaeus 1758) solen ensis ensis siliqua (linnaeus 1758) entalina tetragona (brocchi 1814) enteroxenos oestergreni bonnevie 1902 epitonium clathratulum (kanmacher 1797) epitonium clathrus (linnaeus 1758) scalaria communis epitonium trevelyanum (johnston 1841) epitonium turtonis (turton 1819) scalaria turtonae eulima bilineata (alder 1848) eulima distorta → vitreolina philippii eulimella acicula → eulimella laevis eulimella laevis (brown 1827) syrnola laevis eulimella acicula eulimella nitidissima → ebala nitidissima eulimella scillae (scacchi 1835) evalea divisa → ondina divisa fabulina fabula (gmelin 1791) angulus fabulus tellina fabula fusus antiquus → neptunea antiqua gari depressa (pennant 1777) psammobia vespertina gari fervensis (gmelin 1791) psammobia faeroeensis gari tellinella (lamarck 1818) gastrana fragilis (linnaeus 1758) gibbula cineraria (linnaeus 1758) trochus cineraria gibbula tumida (montagu 1803) trochus tumida glossus humanus (linnaeus 1758) gouldia minima (montagu 1803) circe minima graphis albida (kanmacher 1798) haliella stenostoma (jeffreys 1858) haminea navicula → haminoea navicula haminoea navicula (da costa 1778) haminea navicula hanleya hanleyi (bean 1844) helcion pellucidum (linnaeus 1758) nacella pellucidum patina pellucida hemiaclis ventrosa (jeffreys ms fricle 1874) heteranomia squamula (linnaeus 1758) anomia squamula anomia aculeata hiatella arctica (linnaeus 1758) saxicava arctica hiatella rugosa (linnaeus 1758) saxicava rugosa hinia incrassata (ström 1768) nassa incrassata hinia pygmaea (lamarck 1822) nassa pygmaea hinia reticulata (linnaeus 1758) nassa reticulata omalogyra atomus → omalogyra atomus hydrobia neglecta muus 1963 hydrobia stagnorum → hydrobia ventrosa hydrobia ulvae (pennant 1777) peringia ulvae hydrobia ventrosa (montagu 1803) hydrobia stagnorum iothia fulva (müller 1776) ischnochiton albus (linnaeus 1767) jujubinus clelandi (w. wood 1828) kellia suborbicularis (montagu 1803) kelliella miliaris (philippi 1844) kennerleya glacialis → pandora glacialis geus bulletin no 3.pmd 28-06-2004, 08:46182 183 lacuna crassior (montagu 1803) lacuna divaricata → lacuna vincta lacuna pallidula (da costa 1778) stenotis palidula lacuna parva (montagu 1803) lacuna puteolus lacuna puteolus → lacuna parva lacuna vincta (montagu 1803) lacuna divaricata laevicardium crassum (gmelin 1791) cardium norvegicum lamellaria perspicua (linnaeus 1758) leda minuta → nuculana minuta leda pernula → nuculana pernula lepeta caeca (müller 1776) lepidochitona cinereus (linnaeus 1767) leptaxinus ferruginosus (forbes 1844) axinus ferruginosus leptochiton asellus (gmelin 1791) lepton nitidum (turton 1822) lepton squamosum (montagu 1803) limacina balea → limacina retroversa limacina retroversa (fleming 1823) spiralis retroversus limacina balea spiralis balea limaria hians (gmelin 1791) limaria loscombi (sowerby 1832) limatula subauriculata (montagu 1808) liomesus ovum (turton 1825) liostomia clavula (lovén 1846) litorina rudis → littorina saxatilis litorina rudis → littorina tenebrosa littorina littorea (linnaeus 1758) littorina mariae sacchi & rastelli 1966 littorina obtusata (linnaeus 1758) littorina saxatilis (olivi 1792) litorina rudis littorina tenebrosa (montagu 1803) litorina rudis lucina borealis → lucinoma borealis lucina divaricata → lucinella divaricata lucinella divaricata (linnaeus 1758) lucina divaricata divaricella divaricata lucinoma borealis (linnaeus 1758) lucina borealis lucinopsis undata → mysia undata lunatia alderi (forbes 1838) natica intermedia natica alderi lunatia catena (da costa 1778) natica catena lunatia montagui (forbes 1838) lunatia pallida (broderip & sowerby 1829) natica groenlandica lutraria elliptica → lutraria lutraria lutraria lutraria (linnaeus 1758) lutraria elliptica lymnaea peregra (müller 1774) lyonsia arenosa (möller 1842) lyonsia norvegica (gmelin 1791) macoma balthica (linnaeus 1758) tellina balthica macoma calcarea (gmelin 1791) tellina calcaria macoma loveni jensen 1904 tellina loveni macoma torelli jensen 1904 tellina crassula mactra solida → spisula solida mactra stultorum (linnaeus 1758) mactra corallina cinerea mactra subtruncata → spisula subtruncata mactra corallina cinerea → mactra stultorum malletia obtusa (g.o. sars 1872) mangelia attenuata (montagu 1803) mangelia brachystoma (philippi 1844) mangelia costata → cytharella coarctata mangelia nebula (montagu 1803) margarites helicinus (phipps 1774) melanella alba (da costa 1778) melanella lubrica (monterosato 1891) melaraphe neritoides (linnaeus 1758) menestho divisa → ondina divisa modiola modiolus → modiolus modiolus modiolaria discors → musculus discors modiolaria laevigata → musculus laevigatus modiolaria marmorata → modiolaria tumida modiolaria nigra → musculus niger modiolaria tumida (hanley 1843) musculus tumidus modiolaria marmorata modiolula phaseolina (philippi 1844) mytilus phaseolinus modiolus adriaticus (lamarck 1819) mytilus adriaticus modiolus modiolus (linnaeus 1758) mytilus umbilicatus modiola modiolus montacuta bidentata → mysella bidentata montacuta ferruginosa → tellimya ferruginosa geus bulletin no 3.pmd 28-06-2004, 08:46183 184 montacuta substriata (montagu 1803) musculus discors (linnaeus 1767) modiolaria discors musculus laevigatus (gray 1824) modiolaria laevigata musculus niger (gray 1824) modiolaria nigra musculus tumidus → modiolaria tumida mya arenaria linnaeus 1758 mya truncata linnaeus 1758 myrtea spinifera (montagu 1803) mysella bidentata (montagu 1803) montacuta bidentata mysella dawsoni (jeffreys 1864) mysella tumidula (jeffreys 1867) mysia undata (pennant 1777) lucinopsis undata mytilaster lineatus (gmelin 1791) mytilaster solidus (poli 1795) mytilus minimus mytilus adriaticus → modiolus adriaticus mytilus edulis linnaeus 1758 mytilus minimus → mytilaster solidus mytilus phaseolinus → modiolula phaseolina mytilus umbilicatus → modiolus modiolus nacella pellucidum → helcion pellucidum nassa incrassata → hinia incrassata nassa pygmaea → hinia pygmaea nassa reticulata → hinia reticulata natica affinis (gmelin 1790) natica clausa natica alderi → lunatia alderi natica catena → lunatia catena natica clausa → natica affinis natica groenlandica → lunatia pallida natica intermedia → lunatia alderi neptunea antiqua (linnaeus 1758) fusus antiquus neptunea despecta (linnaeus 1758) neritina fluviatilis → theodoxus fluviatilis nototeredo norvegica (spengler 1792) nucella lapillus (linnaeus 1758) purpura lapillus nucula nitida → nucula nitidosa nucula nitidosa winckworth 1930 nucula nitida nucula nucleus (linnaeus 1767) nucula sulcata (bronn 1831) nucula tenuis → nuculoma tenuis nuculana minuta (müller 1776) leda minuta nuculana pernula (müller 1776) leda pernula nuculoma hanleyi winckworth 1931 nuculoma tenuis (montagu 1808) nucula tenuis obtusella alderi (jeffreys 1858) ocenebra erinacea (linnaeus 1758) odostomia acuta jeffreys 1848 odostomia albella lovén 1846 odostomia unidentata odostomia conoidea winckworth 1932 odostomia eulimoides → brachystomia eulimoides odostomia insculpta → ondina divisa odostomia pallida → brachystomia eulimoides odostomia plicata (montagu 1803) odostomia rissoides → odostomia scalaris odostomia scalaris macgillivray 1843 odostomia rissoides odostomia turrita hanley 1844 odostomia umbilicaris (malm 1863) odostomia unidentata → odostomia albella oenopota incisula (verrill 1882) bela incisula oenopota trevelliana (turton 1834) bela trevelliana oenopota turricola (montagu 1803) bela turricola bela nobilis oenopota violacea (mighels & adams 1842) bela violacea omalogyra atomus (phillippi 1841) omalogyra atomus ondina diaphana (jeffreys 1848) ondina divisa (j. adams 1797) menestho divisa evalea divisa odostomia insculpta onoba aculeus (gould 1841) onoba proxima (forbes & hanley 1850) cingula proxima onoba semicostata (montagu 1803) rissoa striata cingula striata onoba vitrea (montagu 1803) cingula vitrea rissoa vitrea ostrea edulis linnaeus 1758 ovatella myosotis (draparnaud 1801) palliolum greenlandicum (sowerby 1842) pecten groenlandicus palliolum striatum (müller 1776) geus bulletin no 3.pmd 28-06-2004, 08:46184 185 palliolum tigerinum (müller 1776) pandora glacialis leach 1819 kennerleya glacialis panomya arctica (lamarck 1818) paphia aurea (gmelin 1791) tapes aureus paphia aurea senescens (cocconi 1873) tapes senescens tapes aureus parthenia indistincta → chrysallida indistincta parthenia interstincta → chrysallida obtusa parthenia spiralis → chrysallida spiralis parthenina eximia → chrysallida eximia parvicardium exiguum (gmelin 1791) cardium exiguum parvicardium minimum (philippi 1836) cardium minimum parvicardium ovale (sowerby 1840) cardium fasciatum parvicardium scabrum (philippi 1844) cardium nodosum patella vulgata linnaeus 1758 patina pellucida → helcion pellucidum pecten groenlandicus → palliolum greenlandicum pecten islandicus → chlamys islandica pecten maximus (linnaeus 1758) pecten opercularis → aequipecten opercularis pecten septemradius → pseudamussium septemradiatum pecten similis → similipecten similis pecten varius → chlamys varia pelseneeria stylifera (turton 1825) peringia ulvae → hydrobia ulvae phaxas pellucidus (pennant 1777) cultellus pellucidus solen pellucidus philbertia purpurea → raphitoma purpurea philine aperta (linnaeus 1767) philine aperta quadripartita philine catena (montagu 1803) philine denticulata (adams 1800) philine punctata (adams 1800) philine quadrata (s. wood 1839) philine scabra (müller 1776) philine aperta quadripartita → philine aperta philinoglossa helgolandica hertling 1932 pholas candida → barnea candida pholas dactylus linnaeus 1758 plagiocardium papillosum poli 1795 cardium papillosum pododesmus patelliformis (linnaeus 1761) anomia patelliformis pododesmus squama (gmelin 1791) polygireulima monterosatoi (monterosato 1890) polygireulima sinuosa (sacco 1836) portlandia arctica (gray 1824) yoldia arctica portlandia frigida → yoldiella frigida portlandia lenticula → yoldiella lenticula portlandia lucida → yoldiella lucida portlandia tenuis → yoldiella philippiana potamopyrgus antipodarum (gray 1853) psammobia faeroeensis → gari fervensis psammobia vespertina → gari depressa pseudamussium septemradiatum (müller 1776) pecten septemradius psiloteredo megotara (forbes & hanley 1848) puncturella noachina (linnaeus 1771) purpura lapillus → nucella lapillus putilla semistriata → cingula semistriata raphitoma asperrima (brown 1827) raphitoma leufroyi (michaud 1821) raphitoma linearis (montagu 1803) claturella linearis raphitoma purpurea (montagu 1803) philbertia purpurea retusa obtusa (montagu 1803) utriculus obtusus utriculus pertenuis retusa truncatula (bruguière 1792) utriculus truncatulus utriculus mammillatus retusa umbilicata (montagu 1803) utriculus umbilicatus utriculus nitidulus rhizorus acuminatus (bruguière 1792) rissoa abyssicola → alvania abyssicola rissoa albella lovén 1846 turboella albella rissoa cimicoides → alvania cimicoides rissoa inconspicua alder 1844 turboella inconspicua rissoa interrupta → rissoa parva rissoa jan mayeni → alvania jan mayeni rissoa lactea → alvania lactea rissoa lilacina → rissoa violacea rissoa membranacea (j. adams 1800) rissoa parva (da costa 1779) turboella interrupta rissoa interrupta rissoa punctura → alvania punctura rissoa scrobiculata → alvania scrobiculata geus bulletin no 3.pmd 28-06-2004, 08:46185 186 rissoa semistriata → cingula semistriata rissoa striata → onoba semicostata rissoa violacea desmarest 1814 rissoa lilacina rissoa vitrea → onoba vitrea saxicava arctica → hiatella arctica saxicava rugosa → hiatella rugosa saxicavella jeffreysi winckworth 1930 turneria jeffreysi arcinella plicata scalaria communis → epitonium clathrus scalaria turtonae → epitonium turtonis scaphander lignarius (linnaeus 1758) scaphander punctostriatus (mighels & adams 1841) scissurella crispata fleming 1828 scrobicularia piperata → scrobicularia plana scrobicularia plana (da costa 1778) scrobicularia piperata serripes groenlandicus (bruguière 1798) cardium groenlandicum similipecten similis (laskey 1811) pecten similis siphonodentalium lobatum (sowerby 1860) skenea basistriata (jeffreys 1877) skenea planorbis → skeneopsis planorbis skenea serpuloides (montagu 1808) skeneopsis planorbis (fabricius 1780) skenea planorbis solecurtus chamasolen (da costa 1778) solecurtus scopula (turlok 1822) solen ensis → ensis ensis solen pellucidus → phaxas pellucidus spiralis balea → limacina retroversa spiralis retroversus → limacina retroversa spisula elliptica (brown 1827) spisula solida (linnaeus 1758) mactra solida spisula subtruncata (da costa 1778) mactra subtruncata stenotis palidula → lacuna pallidula syndesmya alba → abra alba syndesmya ovata → abra segmentum syndesmya prismatica → abra prismatica syrnola laevis → eulimella laevis tapes aureus → paphia aurea senescens tapes aureus → paphia aurea tapes decussatus (linnaeus 1758) tapes edulis → venerupis rhomboides tapes pullastra → venerupis pullastra tapes senescens → paphia aurea senescens tapes virgineus → venerupis rhomboides taranis borealis bouchet & warén 1980 taranis moerchi (malm 1861) tectura virginea → acmaea virginea tellimya ferruginosa (montagu 1803) montacuta ferruginosa tellina balthica → macoma balthica tellina calcaria → macoma calcarea tellina crassula → macoma torelli tellina donacina linnaeus 1758 tellina fabula → fabulina fabula tellina loveni → macoma loveni tellina pusilla → tellina pygmaea tellina pygmaea (lovén 1846) tellina pusilla tellina tenuis → angulus tenuis teredo navalis linnaeus 1758 theodoxus fluviatilis (linnaeus 1758) neritina fluviatilis thracia convexa (wood 1815) thracia gracilis (jeffreys 1865) thracia papyracea → thracia phaseolina thracia phaseolina (lamarck 1818) thracia papyracea thracia villosiuscula (macgillivray 1827) thyasira croulinensis (jeffreys 1847) thyasira flexuosa (montagu 1803) axinus flexuosus thyasira sarsi (philippi 1845) timoclea ovata (pennant 1777) venus ovata tonicella marmorea (fabricius 1780) tonicella rubra (linnaeus 1767) tridonta borealis schumacher 1817 astarte borealis tridonta elliptica (brown 1827) astarte compressa tridonta montagui (dillwyn 1817) astarte banksie triforis perversa → triphora adversa triforis perversa adversa → triphora adversa triphora adversa (montagu 1803) triforis perversa triforis perversa adversa trivia arctica (pulteney 1799) trivia monacha (da costa 1778) cypraea europaea trochus cineraria → gibbula cineraria trochus tumida → gibbula tumida trophon clathratus → boreotrophon clathratus trophonopsis barvicensis (johnston 1825) troschelia bernicensis (king 1846) geus bulletin no 3.pmd 28-06-2004, 08:46186 187 turboella albella → rissoa albella turboella inconspicua → rissoa inconspicua turboella interrupta → rissoa parva turbonilla acuta → turbonilla delicata turbonilla crenata (brown 1827) turbonilla rufa turbonilla delicata (monterosato 1874) turbonilla acuta turbonilla lactea (linné 1758) chemnitzia lactea turbonilla rufa → turbonilla crenata turbonilla sinuosa (jeffreys 1884) turneria jeffreysi → saxicavella jeffreysi turrisipho moebii (dunker & metzger 1874) turritella communis risso 1826 turritella terebra turritella erosa couthouy 1838 turritella terebra → turritella communis turtonia minuta (fabricius 1780) cyamium minutum utriculus mammillatus → retusa truncatula utriculus nitidulus → retusa umbilicata utriculus obtusus → retusa obtusa utriculus pertenuis → retusa obtusa utriculus truncatulus → retusa truncatula utriculus umbilicatus → retusa umbilicata velutina plicatilis (müller 1776) velutina velutina (müller 1776) venerupis pullastra (montagu 1803) tapes pullastra venerupis rhomboides (pennant 1777) tapes edulis tapes virgineus venus fasciata → clausinella fasciata venus gallina → chamelea striatula venus ovata → timoclea ovata vitreolina collensi (sykes 1903) vitreolina philippii (rayneval & ponzi 1854) eulima distorta balcis devians xylophaga dorsalis turton 1822 yoldia arctica → portlandia arctica yoldia hyperborea lovén 1859 yoldiella frigida (torell 1859) portlandia frigida yoldiella lenticula (möller 1842) portlandia lenticula yoldiella lucida (lovén 1846) portlandia lucida yoldiella philippiana (nyst 1845) portlandia tenuis zirfaea crispata (linnaeus 1758) geus bulletin no 3.pmd 28-06-2004, 08:46187 188 index of species a abra alba 15, 17, 19, 20, 22, 23, 87, 105, 113, 114, 116, 121, 123, 134, 135, 139, 142, 145, 148, 150, 160, 161, 162, 168, 180, 186, 201, 204, 209, 213 abra nitida 19, 20, 22, 23, 87, 111, 123, 124, 129, 142, 145, 148, 150, 156, 157, 180, 201, 205, 209, 213 abra prismatica 19, 22, 23, 87, 88, 106, 113, 114, 119, 121, 123, 142, 148, 150, 166, 180, 186, 201, 205, 209, 213 abra segmentum 88, 113, 114, 121, 172, 180, 186, 205, 213 acanthocardia echinata 17, 18, 20, 22, 23, 78, 105, 113, 114, 118, 119, 120, 121, 123, 139, 141, 143, 145, 148, 150, 180, 181, 201, 204, 208, 213 acanthocardia-venerupis 155 acar nodulosa 66, 180, 200, 208 aclis ascaris 19, 40, 144, 180, 199, 202, 207, 212 aclis minor 18, 19, 22, 41, 104, 120, 141, 144, 150, 167, 180, 199, 202, 207, 212 aclis walleri 19, 41, 144, 180, 199, 202, 207, 212 acmaea tessulata 17, 26, 140, 165, 180, 198, 202, 205, 210 acmaea virginea 17, 21, 26, 140, 147, 162, 165, 180, 186, 198, 202, 206, 211 acteon tornatilis 18, 20, 21, 55, 121, 123, 124, 141, 144, 148, 166, 180, 199, 203, 208, 212 adipicola simpsoni 69, 180, 200, 208 admete viridula 50, 127, 180, 199, 203, 205, 210 aequipecten opercularis 18, 21, 69, 121, 141, 148, 180, 185, 200, 204, 208, 212 akera bullata 15, 17, 18, 21, 59, 113, 114, 121, 134, 138, 141, 148, 162, 165, 180, 200, 203, 208, 212 alvania abyssicola 31, 123, 180, 185, 198, 202, 207, 211 alvania cimicoides 21, 31, 147, 149, 167, 180, 185, 202, 211 alvania cruenta 32, 127, 172, 180, 202, 209 alvania jan mayeni 32, 127, 185, 202, 210 alvania jeffreysi 31, 180, 198, 207 alvania lactea 17, 21, 31, 142, 143, 148, 149, 165, 166, 180, 185, 202, 213 alvania punctura 17, 21, 32, 141, 147, 180, 185, 198, 202, 207, 211 alvania scrobiculata 32, 127, 180, 185, 202, 210 amauropsis islandicus 38, 180, 198, 205 amphilepis norvegica 126, 156 angulus fabulus 182 angulus tenuis 15, 19, 20, 22, 23, 83, 111, 113, 121, 134, 135, 142, 145, 148, 150, 180, 186, 201, 204, 209, 213 anomia aculeata 182 anomia ephippium 18, 22, 72, 142, 148, 180, 204, 213 anomia patelliformis 180, 185 anomia squamula 180, 182 antalis agile 61, 180, 200, 208 antalis entalis 61, 125, 180, 182, 200, 203, 206, 211 antalis eutalis 126 aporrhais pespelicani 16, 18, 19, 21, 22, 36, 108, 113, 116, 120, 123, 135, 136, 141, 144, 147, 150, 160, 162, 165, 180, 198, 202, 207, 212 aporrhais serresianus 36, 180, 198, 207 arca glacialis 180, 181 arca-astarte crenata 133 arcinella plicata 180, 186 arcopagia crassa 83, 180, 201, 209 arctica islandica 10, 15, 17, 19, 20, 22, 23, 88, 101, 111, 112, 114, 116, 120, 122, 130, 131, 132, 134, 135, 138, 140, 143, 144, 147, 149, 154, 172, 180, 182, 201, 205, 206, 211 assiminea grayana 35, 181, 198, 207 astarte banksie 181, 186 astarte borealis 131, 181, 186 astarte compressa 181, 186 astarte sulcata 77, 181, 201, 207 axinopsida orbiculata 74, 127, 128, 130, 181, 204, 210 axinopsis orbiculata 181 axinus ferruginosus 181 axinus flexuosus 181, 186 b balanus balanoides 131 balanus crenata 128 balcis devians 181, 187 barleeia unifasciata 22, 31, 150, 181, 202, 211 geus bulletin no 3.pmd 28-06-2004, 08:46188 189 barnea candida 15, 19, 20, 22, 23, 94, 106, 113, 121, 134, 142, 145, 148, 150, 181, 185, 201, 205, 209, 213 bathyarca glacialis 66, 101, 127, 128, 132, 133, 159, 172, 180, 181, 204, 209 bathyarca pectunculoides 66, 181, 200, 206 bela exarata 48, 122, 181, 199, 203, 205, 210 bela incisula 181 bela nobilis 181 bela trevelliana 181 bela turricola 181 bela violacea 181 bittium reticulatum 16, 17, 18, 19, 21, 22, 35, 107, 113, 118, 120, 123, 124, 129, 134, 135, 136, 138, 141, 144, 147, 150, 157, 160, 162, 164, 171, 181, 198, 202, 207, 212 bittium-varicorbula 155 boreotrophon clathratus 44, 130, 131, 155, 181, 186, 199, 202, 205, 210 boreotrophon truncatus 44, 181, 199, 206 brachystomia carozzai 50, 181, 199, 208 brachystomia eulimoides 17, 18, 19, 21, 50, 121, 138, 141, 144, 147, 181, 184, 199, 203, 208, 212 brissopsis lyrifera 76 buccinum cyaneum 45, 130, 181, 202, 210 buccinum groenlandicum 158, 181 buccinum undatum 17, 18, 19, 21, 22, 45, 108, 118, 119, 120, 123, 131, 138, 140, 144, 147, 150, 162, 172, 181, 199, 202, 206, 211 c cadulus jeffreysi 100, 181 cadulus subfusiforme 60, 125, 126, 181, 200, 203, 208, 212 caecum glabrum 18, 19, 21, 35, 113, 120, 141, 144, 147, 162, 165, 181, 198, 202, 207, 212 calliostoma formosa 27, 181, 198, 207 calliostoma zizyphinum 27, 181, 198, 207 callochiton septemvalvis 25, 181, 198, 207 capulus ungaricus 37, 181, 198, 207 cardium edule 181 cardium groenlandicum 181, 186 cardium minimum 181 cardium nodosum 181 cardium norvegicum 181 cardium papillosum 181, 185 cerastoderma edule 15, 16, 17, 18, 20, 22, 80, 113, 114, 116, 118, 121, 134, 136, 139, 141, 143, 145, 148, 154, 159, 160, 161, 162, 181, 201, 204, 209, 213 cerastoderma glaucum 16, 17, 18, 80, 136, 139, 141, 181, 201, 204, 209, 213 cerithiella metula 39, 181, 199, 207 cerithiopsis barleei 18, 39, 141, 181, 199, 202, 207, 212 cerithiopsis tubercularis 18, 40, 121, 141, 181, 202, 212 chamelea striatula 19, 20, 22, 23, 89, 105, 111, 113, 114, 121, 124, 142, 145, 148, 150, 161, 162, 168, 170, 181, 187, 201, 205, 209, 213 chemnitzia lactea 181, 187 chlamys islandica 69, 128, 130, 131, 158, 181, 185, 204, 211 chlamys varia 18, 20, 22, 23, 70, 106, 121, 141, 145, 148, 150, 162, 181, 185, 200, 204, 208, 212 chrysallida eximia 18, 51, 123, 124, 140, 181, 185, 203, 211 chrysallida decussata 18, 23, 51, 109, 141, 150, 165, 181, 199, 203, 208, 212 chrysallida indistincta 18, 19, 21, 51, 141, 144, 147, 165, 181, 185, 199, 203, 208, 212 chrysallida obtusa 18, 21, 51, 113, 121, 141, 147, 181, 185, 199, 203, 208, 212 chrysallida spiralis 17, 18, 19, 21, 52, 113, 114, 118, 119, 121, 138, 141, 144, 147, 181, 185, 199, 203, 208, 212 cingula semistriata 17, 21, 32, 141, 147, 181, 185, 186, 198, 202, 207, 212 cingula turgida 19, 32, 144, 181, 198, 202, 207, 211 circe minima 182 claturella linearis 185 clausinella fasciata 19, 20, 22, 89, 142, 145, 148, 149, 181, 187, 201, 205, 209, 213 clinocardium ciliatum 80, 122, 124, 127, 154, 156, 181, 204, 210 clione limacina 60, 132, 181, 200, 203, 206, 210 cochlodesma praetenue 20, 23, 96, 106, 125, 145, 150, 170, 181, 201, 205, 209, 213 colus gracilis 45, 181, 199, 208 colus jeffreysianus 45, 181, 199, 208 colus sabini 46, 181, 199, 207 corbula gibba 15, 16, 17, 19, 20, 22, 23, 92, 105, 106, 111, 113, 116, 119, 120, 121, 124, 134, 136, 139, 142, 143, 145, 148, 150, 159, 160, 161, 162, 164, 170, 181, 201, 205, 209, 213 crenella decussata 69, 127, 181, 200, 204, 205, 210 crepidula fornicata 36, 37, 181, 198, 207 cultellus pellucidus 181, 185 geus bulletin no 3.pmd 28-06-2004, 08:46189 190 cuspidaria obesa (lovén 1846) 97 cuspidaria cuspidata 97, 181, 201, 209 cyamium minutum 182, 187 cylichna alba 18, 20, 23, 56, 110, 128, 130, 131, 140, 143, 149, 182, 203, 210 cylichna cylindracea 18, 21, 55, 141, 148, 166, 182, 199, 203, 208, 212 cylichna occulta 56, 126, 127, 130, 131, 158, 172, 182, 203, 209 cylichna propinqua 182 cylichna scalpta 182 cypraea europaea 182, 186 cyprina islandica 116, 152, 180, 182 cytharella coarctata 21, 47, 121, 147, 182, 183, 199, 202, 208, 212 d delectopecten vitreus 18, 70, 100, 125, 126, 140, 156, 165, 182, 200, 204, 206, 211 dentalium entalis 180, 182 dentalium vulgare 62, 99, 125, 172, 182, 203, 213 devonia perrieri 77, 182, 201, 208 diaphana hyalina 182 diaphana minuta 18, 21, 58, 140, 146, 182, 200, 203, 206, 210 divaricella divaricata 182 donax vittatus 19, 20, 21, 22, 23, 85, 86, 106, 110, 111, 121, 122, 142, 143, 145, 148, 150, 155, 167, 169, 170, 182, 201, 204, 209, 213 dosinia exoleta 20, 22, 91, 145, 148, 149, 166, 167, 170, 182, 201, 205, 209, 213 dosinia lincta 20, 22, 91, 113, 114, 121, 145, 148, 149, 166, 167, 169, 170, 182, 201, 205, 209, 213 dosinia lupinus 182 e ebala nitidissima 18, 21, 52, 113, 114, 121, 141, 148, 165, 182, 199, 203, 208, 212 echinocardium cordatum 105, 110 emarginula fissura 26, 182, 198, 207 ensis americanus 82 ensis arcuatus 82, 182, 201, 209 ensis ensis 19, 20, 22, 82, 113, 121, 123, 141, 145, 148, 162, 165, 182, 186, 201, 204, 209, 213 ensis siliqua 83, 182, 201, 209 entalina tetragona 61, 102, 125, 126, 182, 203, 212 enteroxenos oestergreni 44, 182, 199, 207 epitonium clathratulum 40, 182, 199 epitonium clathrus 18, 19, 21, 40, 113, 114, 121, 141, 144, 147, 182, 186, 199, 202, 207, 212 epitonium trevelyanum 22, 40, 150, 182, 199, 202, 207, 212 epitonium turtonis 18, 21, 40, 141, 147, 165, 166, 182, 186, 199, 202, 207, 212 eulima bilineata 41, 182, 199, 207 eulima distorta 182, 187 eulimella acicula 182 eulimella laevis 18, 19, 21, 52, 141, 144, 148, 165, 182, 186, 199, 203, 208, 212 eulimella nitidissima 182 eulimella scillae 18, 23, 52, 104, 105, 123, 124, 141, 150, 167, 182, 199, 203, 208, 212 evalea divisa 182 f fabulina fabula 19, 20, 22, 23, 84, 111, 121, 142, 145, 148, 150, 162, 170, 180, 182, 186, 201, 204, 209, 213 fucus serratus 107 fusus antiquus 182 g gari depressa 22, 86, 148, 149, 167, 182, 185, 204, 213 gari fervensis 19, 22, 23, 86, 111, 140, 147, 149, 150, 166, 182, 185, 201, 204, 207, 211 gari tellinella 86, 182, 201, 209 gastrana fragilis 84, 113, 114, 119, 121, 172, 182, 204, 213 gibbula cineraria 17, 19, 21, 27, 120, 141, 144, 147, 157, 182, 186, 198, 202, 207, 211 gibbula tumida 17, 21, 27, 141, 147, 165, 182, 186, 198, 202, 207, 211 glossus humanus 10, 89, 182, 201, 209 gouldia minima 91, 113, 114, 121, 172, 181, 182, 205, 213 graphis albida 22, 42, 43, 109, 150, 182, 202, 212 h haliella stenostoma 41, 182, 199, 206 haminea navicula 182 haminoea navicula 55, 113, 114, 119, 121, 172, 182, 203, 213 hanleya hanleyi 25, 182, 198, 206 helcion pellucidum 17, 21, 26, 141, 147, 165, 166, 182, 184, 185, 198, 202, 207, 211 geus bulletin no 3.pmd 28-06-2004, 08:46190 191 hemiaclis ventrosa 22, 43, 44, 109, 150, 182, 202, 212 heteranomia squamula 17, 18, 20, 22, 23, 72, 106, 113, 120, 138, 140, 142, 144, 147, 150, 154, 180, 182, 200, 204, 206, 211 hiatella arctica 15, 19, 20, 22, 23, 93, 100, 118, 119, 120, 123, 125, 126, 127, 128, 130, 140, 144, 147, 149, 158, 162, 182, 186, 201, 205, 206, 210 hiatella rugosa 17, 22, 93, 138, 147, 182, 186, 201, 205, 206, 210 hinia incrassata 18, 21, 46, 123, 141, 147, 149, 155, 165, 166, 182, 184, 199, 202, 208, 212 hinia pygmaea 18, 19, 21, 23, 46, 47, 106, 108, 113, 118, 120, 121, 141, 144, 147, 150, 182, 184, 199, 202, 208, 212 hinia reticulata 15, 16, 17, 18, 19, 21, 23, 47, 106, 113, 116, 118, 119, 121, 123, 124, 134, 135, 136, 138, 141, 144, 147, 150, 162, 182, 184, 199, 202, 208, 212 hydrobia neglecta 30, 182, 198, 207 hydrobia ulvae 15, 16, 17, 19, 21, 22, 30, 107, 113, 116, 118, 120, 134, 135, 138, 139, 141, 144, 147, 150, 162, 164, 182, 185, 198, 202, 207, 211 hydrobia ventrosa 15, 16, 17, 19, 30, 31, 134, 135, 138, 141, 144, 198, 202, 207, 211 i iothia fulva 17, 26, 141, 182, 198, 202, 207, 211 ischnochiton albus 25, 182, 198 j jujubinus clelandi 27, 182, 198, 207 k kellia suborbicularis 22, 77, 148, 149, 182, 201, 204, 208, 213 kelliella miliaris 88, 100, 124, 125, 182, 201, 205, 209, 213 l lacuna crassior 29, 183, 198, 205 lacuna pallidula 15, 17, 19, 21, 22, 29, 133, 140, 143, 146, 149, 165, 171, 183, 186, 198, 202, 206, 210 lacuna parva 17, 19, 21, 29, 120, 141, 144, 147, 165, 183, 198, 202, 207, 211 lacuna vincta 16, 17, 19, 21, 30, 120, 123, 128, 131, 134, 135, 136, 138, 140, 144, 147, 162, 183, 198, 202, 206, 211 laevicardium crassum 22, 80, 119, 148, 149, 181, 183, 201, 204, 209, 213 lamellaria perspicua 37, 183, 198, 207 leda minuta 184 leda pernula 133, 184 lepeta caeca 26, 183, 198, 205 lepidochitona cinereus 25, 183, 198, 207 leptaxinus ferruginosus 75, 123, 181, 183, 204, 210 leptochiton asellus 25, 183, 198, 205 lepton nitidum 18, 20, 22, 77, 113, 114, 121, 141, 145, 148, 183, 201, 204, 208, 213 lepton squamosum 77, 183, 201, 208 limacina retroversa 59, 100, 123, 125, 126, 130, 131, 183, 186, 200, 203, 206, 210 limaria hians 73, 183, 200, 208 limaria loscombi 73, 200, 208 limatula subauriculata 73, 183, 200, 206 liomesus ovum 46, 183, 199, 207 liostomia clavula 53, 183, 199, 208 littorina littorea 15, 16, 17, 19, 21, 28, 113, 114, 116, 118, 120, 134, 135, 136, 138, 141, 144, 147, 154, 159, 160, 161, 162, 163, 183, 198, 202, 207, 211 littorina mariae 28, 183, 198, 207 littorina obtusata 15, 16, 17, 19, 21, 28, 29, 120, 134, 138, 139, 140, 144, 147, 162, 183, 198, 202, 206, 211 littorina saxatilis 15, 16, 17, 19, 21, 29, 120, 130, 131, 132, 133, 135, 138, 139, 140, 143, 146, 160, 183, 198, 202, 206, 210 littorina tenebrosa 15, 16, 17, 21, 29, 114, 134, 135, 138, 139, 141, 147, 162, 164, 165, 183, 198, 202, 207, 211 lucinella divaricata 73, 113, 114, 119, 121, 172, 182, 183, 204, 213 lucinoma borealis 18, 22, 73, 141, 143, 148, 149, 155, 166, 183, 200, 204, 208, 213 lunatia alderi 18, 19, 21, 22, 38, 104, 105, 108, 116, 118, 119, 120, 123, 124, 125, 141, 144, 147, 150, 162, 183, 184, 198, 202, 207, 212 lunatia catena 18, 21, 38, 141, 144, 147, 162, 183, 184, 198, 202, 207, 212 lunatia montagui 22, 38, 108, 150, 183, 198, 202, 207, 212 lunatia pallida 39, 127, 128, 130, 131, 183, 184, 198, 202, 205, 210 geus bulletin no 3.pmd 28-06-2004, 08:46191 192 lutraria lutraria 18, 22, 81, 141, 148, 149, 165, 166, 170, 183, 201, 204, 209, 213 lymnaea peregra 15, 16, 60, 134, 135, 163, 183, 200, 203, 208, 212 lyonsia arenosa 96, 127, 130, 172, 183, 205, 210 lyonsia norvegica 23, 96, 150, 183, 201, 205, 209, 213 m macoma balthica 15, 16, 17, 19, 20, 22, 85, 113, 114, 116, 118, 120, 131, 132, 134, 135, 137, 138, 140, 144, 147, 159, 160, 161, 162, 164, 168, 169, 170, 171, 172, 183, 186, 201, 204, 207, 211 macoma calcarea 11, 22, 56, 85, 99, 101, 102, 116, 122, 126, 127, 128, 129, 130, 131, 132, 146, 154, 155, 156, 157, 158, 168, 183, 186, 201, 204, 205, 210 macoma loveni 85, 130, 172, 183, 186, 204, 210 macoma torelli 85, 130, 131, 158, 172, 183, 186, 204, 209 mactra solida 183, 186 mactra stultorum 18, 20, 22, 23, 81, 110, 113, 114, 122, 141, 145, 148, 150, 162, 165, 183, 201, 204, 209, 213 mactra subtruncata 183, 186 malletia obtusa 66, 183, 200, 208 mangelia attenuata 48, 183, 199, 208 mangelia brachystoma 23, 49, 106, 123, 124, 150, 183, 199, 203, 208, 212 mangelia costata 182, 183 mangelia nebula 49, 183, 199, 208 margarites helicinus 17, 27, 140, 164, 183, 198, 202, 205, 210 melanella alba 22, 43, 150, 167, 183, 202, 212 melanella lubrica 22, 43, 104, 150, 183, 199, 202, 207, 212 melaraphe neritoides 28, 183, 198, 207 menestho divisa 183 modiola modiolus 183 modiolaria discors 183, 184 modiolaria laevigata 183 modiolaria marmorata 183 modiolaria nigra 183 modiolaria tumida 18, 21, 69, 121, 141, 148, 183, 200, 204, 208, 212 modiolula phaseolina 15, 18, 21, 67, 113, 114, 121, 134, 135, 141, 148, 183, 184, 200, 204, 208, 212 modiolus adriaticus 18, 21, 67, 141, 148, 166, 183, 184, 200, 204, 208, 212 modiolus modiolus 15, 18, 21, 68, 116, 118, 120, 134, 135, 140, 143, 147, 162, 168, 183, 184, 200, 204, 206, 211 montacuta ferruginosa 114, 186 montacuta substriata 75, 184, 201, 208 musculus discors 15, 17, 18, 20, 21, 23, 68, 110, 133, 138, 140, 144, 146, 149, 183, 184, 200, 204, 206, 210 musculus laevigatus 68, 127, 128, 130, 183, 184, 204, 210 musculus niger 68, 122, 127, 128, 130, 131, 183, 184, 200, 204, 205, 210 musculus tumidus 183 mya arenaria 15, 17, 22, 23, 37, 92, 134, 135, 137, 138, 147, 150, 163, 168, 169, 170, 179, 184, 205, 207, 211 mya truncata 15, 17, 19, 20, 22, 92, 112, 114, 116, 120, 123, 126, 127, 128, 130, 133, 135, 138, 140, 144, 147, 158, 159, 161, 184, 201, 205, 206, 210 myrtea spinifera 74, 184, 201, 208 mysella bidentata 15, 17, 18, 20, 22, 23, 75, 103, 105, 110, 113, 116, 120, 123, 134, 135, 138, 140, 144, 147, 150, 162, 167, 183, 184, 201, 204, 206, 211 mysella dawsoni 76, 184, 201, 206 mysella tumidula 75, 184, 201, 209 mysia undata 19, 22, 91, 121, 142, 148, 183, 184, 201, 205, 209, 213 mytilaster lineatus 67, 119, 121, 122, 172, 184, 204, 213 mytilaster solidus 67, 113, 114, 172, 184, 204, 213 mytilus adriaticus 183 mytilus edulis 15, 16, 17, 18, 20, 21, 23, 51, 67, 68, 106, 110, 112, 114, 116, 118, 119, 120, 123, 124, 129, 131, 132, 134, 135, 137, 138, 139, 140, 143, 144, 147, 150, 156, 157, 158, 159, 160, 161, 162, 172, 184, 200, 204, 206, 211 mytilus phaseolinus 114, 183 mytilus-cerastoderma 155 n nacella pellucidum 182 natica affinis 39, 75, 128, 130, 131, 184, 202, 210 neptunea antiqua 21, 46, 147, 149, 182, 184, 199, 202, 208, 212 neptunea despecta 46, 130, 131, 184, 202, 210 geus bulletin no 3.pmd 28-06-2004, 08:46192 193 neritina fluviatilis 184, 186 nototeredo norvegica 95, 184, 201, 209 nucella lapillus 18, 19, 21, 44, 140, 144, 147, 184, 185, 199, 202, 206, 211 nucula nitida 184 nucula nitidosa 18, 20, 21, 23, 62, 110, 113, 114, 116, 118, 121, 141, 145, 148, 150, 162, 184, 200, 203, 208, 212 nucula nucleus 18, 20, 21, 23, 62, 110, 116, 123, 141, 145, 148, 150, 165, 184, 200, 203, 208, 212 nucula sulcata 20, 63, 113, 114, 121, 123, 124, 145, 184, 200, 204, 208, 212 nucula tenuis 184 nuculana minuta 23, 63, 102, 105, 120, 123, 124, 132, 133, 149, 172, 183, 184, 200, 204, 206, 211 nuculana pernula 64, 99, 100, 102, 122, 125, 126, 127, 129, 130, 131, 132, 155, 183, 184, 200, 204, 205, 210 nuculoma hanleyi 63, 184, 200, 208 nuculoma tenuis 18, 20, 63, 123, 126, 127, 128, 130, 131, 140, 143, 158, 184, 200, 204, 206, 210 o obtusella alderi 32, 184, 198, 207 ocenebra erinacea 44, 184, 199, 209 odostomia acuta 18, 53, 141, 184, 199, 203, 208, 212 odostomia albella 18, 20, 21, 54, 121, 141, 144, 148, 184, 199, 203, 208, 212 odostomia conoidea 15, 18, 19, 21, 23, 53, 134, 135, 141, 144, 148, 150, 184, 199, 203, 208, 212 odostomia eulimoides 184 odostomia insculpta 184 odostomia pallida 184 odostomia plicata 18, 21, 54, 141, 148, 184, 199, 203, 208, 212 odostomia rissoides 184 odostomia scalaris 18, 21, 51, 113, 121, 141, 147, 184, 199, 203, 208, 212 odostomia turrita 18, 21, 54, 123, 141, 148, 184, 199, 203, 208, 212 odostomia umbilicaris 23, 54, 104, 150, 184, 203, 212 odostomia unidentata 184 oenopota incisula 47, 122, 125, 127, 181, 184, 203, 210 oenopota trevelliana 48, 123, 181, 184, 199, 203, 206, 210 oenopota turricola 18, 21, 23, 48, 130, 131, 140, 146, 149, 181, 184, 199, 203, 205, 210 oenopota violacea 48, 122, 181, 184, 203, 210 omalogyra atomus 15, 17, 18, 21, 50, 134, 135, 138, 140, 147, 182, 184, 199, 203, 206, 211 ondina diaphana 18, 19, 53, 141, 144, 165, 184, 199, 203, 208, 212 ondina divisa 18, 52, 123, 141, 165, 182, 183, 184, 199, 203, 208, 212 onoba aculeus 33, 184, 198, 206 onoba proxima 17, 33, 142, 143, 165, 181, 184, 202, 213 onoba semicostata 16, 17, 21, 33, 134, 135, 138, 139, 140, 147, 162, 164, 181, 184, 186, 198, 202, 206, 211 onoba vitrea 16, 17, 19, 21, 22, 33, 104, 105, 120, 138, 141, 144, 147, 150, 167, 181, 184, 186, 198, 202, 207, 212 ostrea edulis 15, 17, 18, 20, 22, 23, 72, 110, 113, 118, 119, 121, 135, 139, 141, 143, 145, 148, 150, 160, 161, 162, 163, 164, 168, 169, 177, 184, 200, 204, 208, 212 ovatella myosotis 60, 184, 200, 208 p palliolum greenlandicum 70, 127, 128, 129, 184, 185, 204, 210 palliolum striatum 18, 71, 141, 165, 184, 200, 204, 208, 212 palliolum tigerinum 18, 71, 141, 165, 185, 200, 204, 208, 212 pandora glacialis 95, 127, 130, 172, 182, 185, 205, 210 panomya arctica 94, 128, 185, 201, 205, 206, 211 paphia aurea 15, 17, 19, 20, 22, 89, 113, 114, 119, 134, 135, 139, 142, 143, 145, 148, 152, 160, 161, 162, 163, 168, 169, 185, 186, 205, 213 paphia senescens 89, 113, 119, 121, 172, 205, 213 parthenia indistincta 181 parthenia interstincta 114, 181 parvicardium exiguum 15, 16, 17, 18, 20, 22, 78, 113, 118, 121, 134, 136, 139, 141, 145, 148, 160, 162, 181, 185, 201, 204, 209, 213 parvicardium minimum 20, 23, 79, 103, 104, 106, 123, 145, 150, 181, 185, 201, 204, 209, 213 parvicardium ovale 15, 18, 20, 22, 79, 120, 123, 134, 135, 140, 144, 147, 169, 181, 185, 201, 204, 207, 211 geus bulletin no 3.pmd 28-06-2004, 08:46193 194 parvicardium scabrum 15, 17, 18, 20, 22, 79, 121, 134, 135, 139, 141, 145, 148, 162, 181, 185, 201, 204, 209, 213 patella vulgata 17, 26, 140, 185, 198, 202, 207, 211 patina pellucida 182 pecten islandicus 181 pecten maximus 22, 71, 148, 149, 167, 185, 200, 204, 208, 212 pecten septemradius 185 pecten similis 185, 186 pecten varius 181, 185 pecten vitreus 126, 156 pelseneeria stylifera 44, 185, 199, 207 peringia ulvae 185 phaxas pellucidus 19, 20, 22, 23, 83, 105, 110, 114, 121, 123, 141, 145, 148, 150, 159, 170, 181, 185, 186, 201, 204, 209, 213 philbertia purpurea 185 philine aperta 18, 21, 56, 113, 121, 141, 148, 185, 199, 203, 208, 212 philine catena 57, 125, 185, 199, 203, 208, 212 philine denticulata 57, 185, 199, 208 philine punctata 18, 21, 57, 141, 148, 185, 199, 203, 208, 212 philine quadrata 57, 185, 199, 206 philine scabra 57, 185, 199, 208 philinoglossa helgolandica 58, 185, 199, 208 pholas candida 181, 185 pholas dactylus 19, 22, 23, 94, 95, 111, 142, 148, 150, 165, 166, 185, 201, 205, 209, 213 plagiocardium papillosum 79, 121, 122, 155, 172, 181, 185, 204, 213 pododesmus patelliformis 18, 22, 71, 141, 148, 180, 185, 200, 204, 208, 212 pododesmus squama 71, 185, 200, 208 polygireulima monterosatoi 42, 185, 199, 207 polygireulima sinuosa 22, 41, 106, 150, 185, 199, 202, 207, 212 portlandia arctica 11, 50, 64, 100, 101, 120, 126, 127, 128, 129, 130, 131, 132, 133, 154, 156, 157, 158, 159, 172, 185, 187, 204, 209 portlandia frigida 65, 66, 133, 185, 187 portlandia lenticula 133, 185, 187 portlandia lucida 185, 187 portlandia tenuis 185, 187 potamopyrgus antipodarum 30, 185, 198, 207 psammobia faeroeensis 182, 185 psammobia vespertina 182, 185 pseudamussium septemradiatum 71, 73, 123, 124, 129, 185, 200, 204, 208, 212 psiloteredo megotara 95, 185, 201, 205 puncturella noachina 27, 155, 185, 198, 206 purpura lapillus 184, 185 putilla semistriata 185 r raphitoma asperrima 49, 185, 199, 208 raphitoma leufroyi 49, 185, 199, 208 raphitoma linearis 18, 21, 49, 123, 124, 141, 147, 165, 181, 185, 199, 203, 208, 212 raphitoma purpurea 18, 49, 141, 165, 185, 203, 212 retusa obtusa 15, 17, 18, 20, 21, 58, 127, 130, 133, 138, 140, 143, 146, 164, 165, 185, 187, 200, 203, 205, 210 retusa truncatula 15, 16, 17, 18, 20, 21, 23, 58, 105, 113, 121, 134, 135, 136, 138, 141, 144, 148, 150, 162, 185, 187, 200, 203, 212 retusa umbilicata 18, 20, 21, 23, 59, 113, 114, 123, 124, 141, 144, 148, 150, 185, 187, 200, 203, 208, 212 rhizorus acuminatus 59, 185, 200, 208 rissoa abyssicola 185 rissoa albella 16, 17, 19, 21, 22, 33, 34, 120, 134, 135, 136, 138, 141, 144, 147, 150, 160, 185, 187, 198, 202, 207, 212 rissoa cimicoides 185 rissoa inconspicua 16, 17, 19, 21, 34, 113, 114, 118, 119, 120, 134, 135, 136, 138, 141, 144, 147, 162, 185, 187, 198, 202, 207, 212 rissoa interrupta 185 rissoa jan mayeni 185 rissoa lactea 185 rissoa lilacina 185, 186 rissoa membranacea 16, 17, 19, 21, 34, 113, 114, 120, 134, 135, 136, 138, 141, 144, 147, 162, 163, 185, 198, 202, 207, 212 rissoa parva 17, 21, 34, 113, 114, 120, 123, 124, 141, 147, 162, 185, 187, 198, 202, 207, 212 rissoa punctura 185 rissoa scrobiculata 185 rissoa semistriata 186 rissoa striata 186 rissoa violacea 17, 19, 21, 22, 34, 107, 118, 119, 120, 141, 144, 147, 150, 165, 185, 186, 198, 202, 207, 212 rissoa vitrea 186 geus bulletin no 3.pmd 28-06-2004, 08:46194 195 s saxicava arctica 133, 186 saxicava rugosa 186 saxicavella jeffreysi 19, 20, 22, 23, 93, 111, 121, 142, 145, 148, 150, 180, 186, 187, 201, 205, 209, 213 scalaria communis 186 scalaria turtonae 186 scaphander lignarius 56, 186, 199, 208 scaphander punctostriatus 56, 186, 199, 206 scissurella crispata 25, 28, 118, 120, 186, 202, 211 scrobicularia piperata 186 scrobicularia plana 15, 16, 17, 19, 20, 22, 86, 113, 114, 116, 119, 121, 134, 136, 139, 142, 143, 145, 148, 160, 162, 163, 186, 201, 204, 209, 213 serripes groenlandicus 80, 122, 127, 156, 181, 186, 204, 210 similipecten similis 71, 123, 185, 186, 200, 204, 208, 212 siphonodentalium lobatum 61, 100, 102, 125, 126, 132, 186, 203, 210 skenea basistriata 17, 28, 141, 165, 186, 198, 202, 207, 211 skenea planorbis 186 skenea serpuloides 17, 28, 142, 186, 202, 213 skeneopsis planorbis 16, 17, 21, 31, 138, 140, 147, 186, 202, 211 solecurtus chamasolen 82, 186, 201 solecurtus scopula 82, 186, 201, 209 solen ensis 182, 186 solen pellucidus 185, 186 spiralis balea 186 spiralis retroversus 186 spisula elliptica 18, 20, 56, 81, 113, 114, 140, 144, 147, 186, 201, 204, 207, 211 spisula solida 18, 20, 22, 81, 119, 141, 145, 148, 183, 186, 201, 204, 209, 213 spisula subtruncata 15, 19, 20, 22, 23, 81, 82, 104, 107, 110, 113, 116, 118, 119, 121, 134, 141, 143, 145, 148, 150, 161, 162, 163, 170, 183, 186, 201, 204, 209, 213 stenotis palidula 186 syndesmya alba 186 syndosmya prismatica 114 syrnola laevis 182 t tapes decussatus 15, 17, 19, 20, 22, 90, 121, 134, 135, 139, 142, 143, 145, 148, 160, 161, 163, 164, 168, 186, 205, 213 tapes edulis 186, 187 tapes pullastra 186, 187 tapes senescens 186 tapes virgineus 186, 187 taranis borealis 50, 186, 199, 207 taranis moerchi 50, 186, 199, 208 tectura virginea 186 tellimya ferruginosa 18, 20, 22, 23, 76, 105, 110, 113, 114, 120, 140, 144, 147, 150, 162, 167, 183, 186, 201, 204, 206, 211 tellina balthica 186 tellina calcaria 186 tellina crassula 186 tellina donacina 84, 121, 186, 204, 213 tellina fabula 182, 186 tellina loveni 186 tellina pusilla 186 tellina pygmaea 23, 84, 111, 150, 186, 201, 204, 213 tellina tenuis 180, 186 teredo navalis 95, 186, 201, 209 theodoxus fluviatilis 16, 19, 28, 138, 139, 144, 184, 186, 198, 202, 207, 211 thracia convexa 96, 186, 201, 209 thracia gracilis 97, 186, 201, 209 thracia papyracea 186 thracia phaseolina 19, 20, 22, 23, 96, 97, 113, 121, 142, 145, 148, 150, 162, 186, 201, 205, 209, 213 thracia villosiuscula 97, 121, 186, 201, 205, 209, 213 thyasira croulinensis 74, 186, 201, 206 thyasira flexuosa 18, 20, 22, 23, 74, 120, 130, 131, 140, 144, 147, 149, 181, 186, 201, 204, 206, 210 thyasira sarsi 75, 186, 201, 205 timoclea ovata 19, 20, 22, 23, 90, 111, 113, 114, 119, 121, 142, 145, 148, 150, 162, 186, 187, 201, 205, 209, 213 tonicella marmorea 25, 131, 172, 186, 198, 202, 206, 211 tonicella rubra 25, 186, 198, 207 tridonta borealis 15, 18, 77, 116, 127, 128, 130, 131, 133, 135, 138, 140, 143, 168, 181, 186, 201, 204, 205, 210 geus bulletin no 3.pmd 28-06-2004, 08:46195 196 tridonta elliptica 78, 116, 120, 122, 127, 138, 168, 181, 186, 201, 204, 205, 210 tridonta montagui 78, 116, 120, 128, 181, 186, 201, 204, 206, 210 triforis perversa 186 triforis perversa adversa 186 triphora adversa 15, 17, 18, 19, 21, 39, 113, 121, 134, 138, 141, 144, 147, 162, 186, 199, 202, 212 trivia arctica 37, 186, 198, 207 trivia monacha 21, 37, 148, 149, 167, 182, 186, 202, 213 trochus cineraria 182 trochus tumida 182 trophonopsis barvicensis 45, 186, 199, 208 troschelia bernicensis 47, 186, 199, 208 turboella albella 185, 187 turboella inconspicua 185, 187 turboella interrupta 185, 187 turbonilla acuta 187 turbonilla crenata 18, 20, 54, 113, 114, 121, 141, 144, 187, 199, 203, 208, 212 turbonilla delicata 18, 20, 23, 55, 141, 144, 150, 187, 199, 203, 208, 212 turbonilla lactea 18, 20, 21, 55, 113, 114, 118, 121, 141, 144, 148, 181, 187, 199, 203, 208, 212 turbonilla rufa 114, 187 turbonilla sinuosa 23, 150, 187, 203, 212 turneria jeffreysi 186, 187 turrisipho moebii 46, 187, 199, 207 turritella communis 18, 19, 21, 22, 35, 36, 105, 116, 118, 119, 120, 123, 124, 125, 139, 141, 144, 147, 150, 151, 154, 156, 168, 187, 198, 202, 207, 212 turritella erosa 36, 124, 127, 128, 156, 157, 172, 187, 202, 209 turritella terebra 100, 154, 156, 157, 187 turtonia minuta 18, 22, 23, 76, 106, 140, 147, 150, 182, 187, 204, 211 u utriculus mammillatus 185, 187 utriculus nitidulus 185, 187 utriculus obtusus 185, 187 utriculus pertenuis 185, 187 utriculus truncatulus 185, 187 utriculus umbilicatus 185, 187 v velutina plicatilis 37, 187, 198, 206 velutina velutina 37, 187, 198, 206 venerupis pullastra 15, 17, 19, 20, 22, 90, 91, 121, 134, 135, 139, 142, 143, 145, 148, 160, 161, 162, 163, 169, 186, 187, 201, 205, 209, 213 venerupis rhomboides 19, 22, 90, 142, 148, 149, 161, 166, 186, 187, 201, 205, 209, 213 venus fasciata 187 venus gallina 114, 181, 187 venus ovata 186, 187 vitreolina collensi 22, 42, 151, 167, 171, 187, 202, 213 vitreolina philippii 18, 21, 22, 42, 109, 123, 141, 147, 150, 165, 181, 182, 187, 199, 202, 207, 212 x xylophaga dorsalis 95, 187, 201, 209 y yoldia arctica 185, 187 yoldia hyperborea 56, 64, 100, 101, 127, 129, 132, 133, 187, 204, 210 yoldiella frigida 65, 100, 123, 125, 126, 128, 132, 185, 187, 204, 210 yoldiella lenticula 65, 102, 122, 127, 130, 131, 132, 185, 187, 204, 210 yoldiella lucida 65, 123, 185, 187, 200, 204, 206, 211 yoldiella nana 65, 66 yoldiella philippiana 66, 123, 185, 187, 204, 212 z zirfaea crispata 15, 19, 20, 22, 94, 112, 114, 122, 129, 130, 131, 132, 134, 140, 144, 147, 172, 187, 201, 205, 206, 211 geus bulletin no 3.pmd 28-06-2004, 08:46196 d ep th b el ow s ur fa ce , m d ep th b el ow s ur fa ce , m d ep th b el ow s ur fa ce , m d ep th b el ow s ur fa ce , m d ep th b el ow s ur fa ce , m 0 10 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 0 10 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 0 10 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 0 10 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 0 10 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 0 25 50 75 100 0 25 50 75 100 0 25 50 75 100 0 25 50 75 100 0 25 50 75 100 0 25 50 75 100 0 25 50 75 100 0 25 50 75 100 0 1 2 3 4 0.0 0.2 0.4 0.6 0.8 0.0 0.1 0.2 0.3 0.4 0.00 0.10 0.20 0.30 0.00 0.10 0.20 0.00 0.05 0.10 0.15 0.20 0.0 0.1 0.2 0.3 0.4 1 2 3 4 5 6 7 8 9 0.0 0.4 0.8 1.2 1.6 2.0 0.0 0.1 0.2 0.3 0.4 0 10 20 30 40 50 60 70 0 10 20 30 40 0 1 2 3 4 5 6 7 8 9 0 10 20 30 0 1000 2000 3000 0.00 0.01 0.02 0.03 0.04 0.00 0.10 0.20 pyrite reworked spatangoids cirripeds ophiuroids concretions fish other fossilsmean grain size q1+q2+q3/3, mm sorting coefficient (so)= q1/q3 geometric skewness q1·q3/q2·q2 kurtosis q1-q3/2(p10-p90) uniformity coefficient p40/p90 water (ds) weight per cent, % loss on ignition (550 c) weight per cent, % number of species number of specimens number of species/ total weight number of specimens/ total weight p10 percentile (10%), mm q1 quartile (25%), mm p40 percentile (40%), mm q2 quartile (50%), mm q3 quartile (75%), mm p90 percentile (90%), mm clay < 0.002 mm weight per cent, % fine silt > 0.002 mm weight per cent, % medium silt > 0.006 mm weight per cent, % coarse silt > 0.02 mm weight per cent, % fine sand > 0.063 mm weight per cent, % medium sand > 0.2 mm weight per cent, % coarse sand > 0.6 mm weight per cent, % gravel > 2.0 mm weight per cent, % a: grain size distribution b: quartiles and percentiles c: classification of fragmental deposits d: species and specimens diversity e: occurrences of minerals, fossils and reworked molluscs fig. 93. the skagen well data other than the molluscan record kennedy et al. 2024: geus bulletin 57. 8366. https://doi.org/10.34194/geusb.v57.8366 1 of 29 research article early cenomanian ammonites from east and north-east greenland w. james kennedy1, simon r. a. kelly2†, simon schneider*2 1oxford university museum of natural history, oxford, uk; 2casp, cambridge, uk †deceased 19 may 2023 abstract early cenomanian (100.5–95.7 ma) ammonite faunas from east and north-east greenland collected by the late simon kelly and colleagues are described. the assemblages are dominated by typically boreal schloenbachia varians (j. sowerby 1817). also present are parapuzosia (austiniceras) austeni (sharpe 1855) and species of more typically tethyan genera. these include phylloceras (hypophylloceras) lombardense (joly 2000), gaudryceras (gaudryceras) cassisianum (d’orbigny 1850), gaudryceras (mesogaudryceras) leptonema (sharpe 1855), and the hypermorphic tetragonitine titanoleioceras boreale gen. et sp. nov. previously known only from geographical society ø and traill ø, the newly described material extends the distribution of these early cenomanian faunas northwards to hold with hope and south to the kangerlussuaq basin. the phylloceratids, gaudryceratids, and tetragonitids in these assemblages are probably not preserved in their preferred original habitats, but rather drifted to their respective sites of burial during or after their lifetime. *correspondence: simon.schneider@ casp.org.uk received: 16 jan 2024 revised: 31 may 2024 accepted: 17 jun 2024 published: 22 aug 2024 keywords: late cretaceous, boreal realm, tethys, taxonomy, palaeobiogeography abbreviations: a: adventive lobe d: diameter e: external lobe gm: geological museum (registration prefix for specimens stored at natural history museum of denmark) i: internal lobe l: lateral lobe mguh: museum geologicum universitatis hafniensis (registration prefix for type and figured specimens stored at natural history museum of denmark) u: umbilical lobe/umbilicus wb: whorl breadth wh: whorl height wgs84: world geodetic system 1984 geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: karen dybkjær (geus, denmark) reviewed by: peter alsen (geus, denmark) and one anonymous reviewer. funding: see page 28 competing interests: see page 28 additional files: see page 28 1 introduction 1.1 rationale and scope in this contribution, we describe early cenomanian (100.5–95.7 ma) ammonites from east and north-east greenland, most of which were collected by the late simon kelly during his time at casp, uk. the presence of cenomanian ammonite faunas in north-east greenland was acknowledged by leonard f. spath (1946), who recognised the classic boreal ammonite schloenbachia amongst material collected by hans stauber on geographical society ø and traill ø. desmond donovan, working in the same area, discovered additional material, documented and illustrated in a series of papers (donovan 1949, 1953, 1954, 1955; summarised in donovan 1957). the material described in this study extends the distribution of these faunas northwards to hold with hope and south to the kangerlussuaq basin (kangersertuaq region; fig. 1) and documents a greater diversity of taxa than previously recorded. simon r.a. kelly passed away in may 2023, prior to the completion of this publication. besides collecting most of the specimens, he had cleaned and prepared the material and made initial taxonomic and biostratigraphic assessments. recognising the scientific value of the ammonites, he initiated the study and had seen and approved the photographic work and large parts of the text before his passing. as previously agreed, simon schneider joined as a co-author to complete simon kelly’s tasks. w. james kennedy is chiefly responsible for the taxonomic part of the study, and the new taxa take his authorship only. https://doi.org/10.34194/geusb.v57.8366 https://orcid.org/0000-0001-6493-357x mailto:simon.schneider@casp.org.uk mailto:simon.schneider@casp.org.uk https://creativecommons.org/licenses/by/4.0/deed.en kennedy et al. 2024: geus bulletin 57. 8366. https://doi.org/10.34194/geusb.v57.8366 2 of 29 geusbulletin.org –18°w–20°w–22°w–24°w–26°w–28°w–30°w–32°w–34°w 74°n 72°n 70°n 68°n jameson land jameson land land liverpoolliverpool land kangertittivaq traill øtraill ø geographical society ø hold with hope hold with hope gåselandgåseland wollaston forland wollaston forland kangersertuaq region (kangerlussuaq basin) blos se vil le kys t blos se vil le kys t kong oscar fjord kangerluk kong oscar ymer øymer ø clavering øclavering ø kuhn ø sc hu ch er t d al sc hu ch er t d al r ød ef jo rd r ød ef jo rd ugleelvugleelv ella øella ø g re en la nd cenozoic basalt cretaceous jurassic triassic permian carboniferous devonian other rock units cenozoic intrusions faults 100 km ‘windy valley’ lygnaelv tværdal leitch bjerg lysdal ilimananngip nunaa ilimananngip nunaa fig. 1 geological overview of the study area in east and north-east greenland. localities where cenomanian ammonites were collected are indicated by yellow stars. from north to south, these are lygnaelv (casp locality no. k7359, coordinates 74.2639, –20.55, wgs84 datum); tværdal (casp locality no. k7350, coordinates 72.9633, –23.0667; no. k7351, coordinates 73.0003, –22.9717; no. w3099, coordinates 72.958, –23.0452); lysdal (casp locality no. w4389, coordinates 72.8957, –22.807; no. w4390, coordinates 72.8965, –22.8122); leitch bjerg (casp locality no. w4342, coordinates 72.8555, –22.5023); and ‘windy valley’ (casp locality no. w4225; coordinates: 68.6383, –30.9297). https://doi.org/10.34194/geusb.v57.8366 http://www.geusbulletin.org/ kennedy et al. 2024: geus bulletin 57. 8366. https://doi.org/10.34194/geusb.v57.8366 3 of 29 geusbulletin.org 2 materials and locality details the specimens described here were collected during several field seasons of casp, between 1994 and 2003, and come from three regions in east and north-east greenland (figs 1, 2). farthest to the north, in the hold with hope region, cenomanian ammonites were found in a single locality at lygnaelv (casp locality no. k7359, coordinates in decimal degrees 74.2639, –20.55, world geodetic system 1984 (wgs84) datum; supplementary file s1; ‘hh’ in table 1), exposing the fosdalen formation (home forland group; kelly et al. 1998; see bjerager et al. 2020 for updated lithostratigraphy). the fosdalen formation is an over 1-km thick unit of dark mudstones with siderite concretions, intercalated with varying proportions of thin fine-grained sandstone beds (kelly et al. 1998), and was deposited in slope to basin floor settings according to bjerager et al. (2020). combined biostratigraphy of dinoflagellate cysts, inoceramid bivalves, and locally ammonites, indicates a middle albian to middle coniacian age (between 108 and 87 ma) for the fosdalen formation on hold with hope (kelly et al. 1998; bjerager et al. 2020). of three fragmentary cenomanian ammonites collected, two are too poorly preserved for identification; the third specimen is assigned to schloenbachia varians forma ventriosa below. hold with hope (kelly et al. 1998; bjerager et al. 2020) h om e fo rla nd g ro up østersletten formation knudshoved formation nanok member fosdalen formation rødelv member stratumbjerg formationb ro rs on h al vø g p langsiden mb geographical society ø (parsons et al. 2017; bjerager et al. 2020) h om e fo rla nd g ro up inoceramus lamarcki beds sphenoceramus beds knudshoved formation inoceramus crippsi beds inoceramus angelicus beds leitch bjerg fm. månedal formation fosdalen formation scaphites beds ? ? ? k an ge rd lu gs su aq g ro up kangerlussuaq basin (larsen et al. 2005; nøhr-hansen 2012) suunigajik member watkins fjord formation sorgenfri formation christian iv formation marine mudstone unconformitymarine conglomerate marine sandstonefluvial sandstone maastrichtian campanian santonian coniacian turonian cenomanian u pp er lo w er c re ta ce ou s albian fig. 2 albian to maastrichtian lithostratigraphy of the kangerlussuaq basin (data sources: larsen et al. 2005; nøhr-hansen 2012), geographical society ø (data sources: parsons et al. 2017; bjerger et al. 2020) and hold with hope (data sources: kelly et al. 1998; bjerger et al. 2020). table 1 occurrences of ammonite species in the eight localities studied marked by ‘×’. species name hh t1 t2 t3 lb ld1 ld2 wv gaudryceras (gaudryceras) cassissianum (d’orbigny 1850) × × gaudryceras (mesogaudryceras) leptonema (sharpe 1855) × × parapuzosia (austiniceras) austeni sharpe 1855 × phylloceras (hypophylloceras) lombardensis (joly 2000) × schloenbachia varians (j. sowerby 1817) × × × × × × × titanoleioceras boreale n. gen., n. sp. × t1–t3: tværdal 1–3. lb: leitch bjerg. ld1, ld2: lysdal 1, 2. hh: lygnaelv, hold with hope. wv: ‘windy valley’. https://doi.org/10.34194/geusb.v57.8366 http://www.geusbulletin.org/ kennedy et al. 2024: geus bulletin 57. 8366. https://doi.org/10.34194/geusb.v57.8366 4 of 29 geusbulletin.org approximately 100 km farther south, six localities on geographical society ø also exposing the fosdalen formation, yielded ammonites (see bjerager et al. 2020 for a revised lithostratigraphy). the fosdalen formation on geographical society ø is thought to be of similar thickness as on hold with hope (cf. bjerager et al. 2020) and represents slope to basin-floor deposits (parsons et  al. 2017). note that parsons et al. (2017, their supplementary materials 4) reported macrofauna including markers of early and late albian, early and middle cenomanian, late turonian and coniacian age from the fosdalen formation, supporting a chronostratigraphic range between 112 and 86 ma. for the overlying knudshoved formation, sensu bjerager et al. (2020), parsons et  al. (2017) reported late santonian to early campanian inoceramids and middle to late campanian ammonites (indicating an age range between 84 and 70 ma). furthermore, parsons et al. (2017) regarded the contact between the fosdalen and knudshoved formations on geographical society ø as conformable. in contrast, bjerager et al. (2020) indicated a major hiatus encompassing the entire coniacian and santonian stages, but did not describe the contact of the two units on geographical society ø. thus, while the precise relationship between these two units remains unclear, the age restrictions imposed by bjerager et al. (2020) are clearly contradicted by macrofossil evidence (parsons et al. 2017). the lithostratigraphic column depicted in fig. 2 represents a merger of both schemes. cenomanian ammonites were collected from three localities in tværdal (fig. 1; table 1: ‘t1’, casp locality no. k7350, coordinates 72.9633, –23.0667; ‘t2’, casp locality no. k7351, coordinates 73.0003, –22.9717; ‘t3’, casp locality no. w3099, coordinates 72.958, –23.0452), two localities in lysdal (fig. 1; table 1: ‘ld1’, casp locality no. w4389, coordinates 72.8957, –22.807; ‘ld2’, casp locality no. w4390, coordinates 72.8965, –22.8122), and a single outcrop on leitch bjerg (fig. 1; table 1: ‘lb’, casp locality no. w4342, coordinates 72.8555, –22.5023). all but one of these localities yielded specimens of schloenbachia varians. in addition, gaudryceras (gaudryceras) cassisianum and gaudryceras (mesogaudryceras) leptonema each occur at a single locality (table 1). another 500 km farther south, in the kangerlussuaq basin, ammonites occur in the sorgenfri formation, an up to c. 140 m thick unit of sandy mudstones with characteristic dark, phosphatic concretions, rare discrete sandstone beds, and occasional calcareous concretions and concretionary horizons (larsen et  al. 2005). sedimentology suggests deposition in a midto outer shelf setting (larsen et  al. 2005). a middle albian to middle coniacian age (between 108 and 87 ma) is indicated by combined ammonite and dinoflagellate cyst biostratigraphy (larsen et  al. 2005). ammonites were collected from a single outcrop in ‘windy valley’ (casp locality no. w4225, coordinates: 68.6383, –30.9297 wgs84 datum), where the sorgenfri formation is more than 95 m thick (fig. 3), consisting of mudstone with sparse bioturbation throughout, and concretions occuring in the lower 30 m of the succession. this locality has produced the richest and most diverse ammonite assemblage of cenomanian age, including all six species described herein. ammonites occur scattered within the lower 60 m of the exposure (fig. 3), but most specimens were collected loose from the base of the outcrop. based on the few specimens collected in situ, there is no apparent zonation. ‘windy valley’ is the type locality of titanoleioceras boreale gen. et sp. nov. generally, it should be noted that only the best-preserved, most representative specimens are included in this study. numerous additional cenomanian ammonite specimens collected by casp are either too poorly preserved to be diagnostic or are more fragmentary representatives of the taxa described below (particularly schloenbachia). all specimens from greenland that are figured and discussed in this study are curated at the natural history museum of denmark, copenhagen, denmark, under accession numbers mguh 34448–34490 and gm 2024.2–2024.14. this published work and the  nomenclatural acts it contains have been registered in zoobank: http://zoobank.org/urn:lsid:zoobank. org:pub:4f83504a-7c33-418d-b812-04dc558f16e4 3 discussion: age and composition of the faunas the widespread occurrence of the genus schloenbachia, identified as the early cenomanian schloenbachia varians (j. sowerby 1817), confirms the horizon of the ammonites described below. this species dominates the overall assemblage and is a classic boreal indicator, extending from north-east greenland eastwards on the north side of tethys to iran north of the zagros and western kazakhstan (fig. 4; wilmsen & mosavinia 2010, fig. 2; kennedy 2013, text-fig. 1). records from algeria on the south side of tethys (kennedy & juignet 1984, p. 123; mendir et al. 2019, p. 239, plate 1, fig. 3) are based on poorly preserved acompsoceras hyatt 1903, in our view. of other species present, parapuzosia (austiniceras) austeni (sharpe 1855) was previously known from the cenomanian and turonian, ranging from southern england to france, germany, the czech republic, ukraine (crimea), and, perhaps, kwazulu-natal in south africa (wright & kennedy 1984, p. 60). most remarkable is the diversity, albeit in small numbers of specimens, of what are generally regarded as classic tethyan genera that are rare in the boreal faunas of europe: phylloceras (hypophylloceras) salfeld 1924, gaudryceras (gaudryceras) de grossouvre 1894, and g. (mesogaudryceras) spath 1927. the giant titanoleioceras gen nov. is believed to have its origins in the typically tethyan tetragonitinae. https://doi.org/10.34194/geusb.v57.8366 http://www.geusbulletin.org/ http://zoobank.org/urn:lsid:zoobank.org:pub:4f83504a-7c33-418d-b812-04dc558f16e4 http://zoobank.org/urn:lsid:zoobank.org:pub:4f83504a-7c33-418d-b812-04dc558f16e4 kennedy et al. 2024: geus bulletin 57. 8366. https://doi.org/10.34194/geusb.v57.8366 5 of 29 geusbulletin.org s 0 10 60 7020 30 40 50 80 90 94.5 m mudstone concreted horizon pebbly sandstone dolerite sill ammonite concretion bioturbation m t t a a t c b a fig. 3 photographs and schematic log of casp locality no. w4225, ‘windy valley’, kangerlussuaq basin. a: upper half of exposure, with prominent, yellowish pebbly sandstone bed at 67 m log height (person for scale). b: close-up view along pebbly sandstone bed at 67 m log height (1.5 m walking pole for scale). c: ammonite parapuzosia (austiniceras) austeni (sharpe 1855) in situ in fissile shale at approximately 12 m log height (mguh 34452; pencil for scale). original log and all photographs by a.g. whitham (casp). abbreviations: a: parapuzosia (austiniceras) austeni. m: gaudryceras (mesogaudryceras) leptonema. s: schloenbachia varians. t: titanoleioceras boreale. https://doi.org/10.34194/geusb.v57.8366 http://www.geusbulletin.org/ kennedy et al. 2024: geus bulletin 57. 8366. https://doi.org/10.34194/geusb.v57.8366 6 of 29 geusbulletin.org in contrast, cosmopolitan members of the acanthoceratoidea, which provide zonal indices for the cenomanian of the eurasian part of the boreal realm, are absent, as are heteromorphs. in lower cenomanian successions, such as those in the west melbury formation marly chalks of south-east england or the phosphatic faunas of the glauconitic marl in the isle of wight, specimens of schloenbachia varians make up perhaps as much as 99% of the ammonite fauna. absence of acanthoceratoidea and heteromorphs from the much smaller collections from east and northeast greenland may thus simply reflect collection failure. the phylloceratid, gaudryceratid, and tetragonitid occurrences are unlikely to be a record of populations in their optimum habitats. rather, they probably represented individuals that became separated from their parent populations, and drifted, either in life or after death, to their final site of burial. quite where those parent populations were located is unknown. dispersal of ammonites over long distances is documented and discussed by kennedy & cobban (1976, p. 64; see in particular text-fig. 19). a recent example of such an occurrence even more remarkable than the present ones is the single gaudryceratid recorded from the late cenomanian and earliest turonian hydrothermal vent faunas of the troodos massif in cyprus that formed at depths of 2500–5000 m on arc-related spreading ridges of the neotethys (kaim et  al. 2021, p. 1942, text-fig. 5f), and interpreted as a waterlogged shell that sank from surface waters. 4 systematic palaeontology conventions. dimensions are given in mm: d: diameter. wb: whorl breadth. wh: whorl height. u: umbilicus. figures given in parentheses are dimensions as a percentage of diameter. the suture terminology is that of korn et  al. (2003): e: external lobe. a: adventive lobe. u1, u2 and u2: umbilical lobes. i: internal lobe. l: lateral lobe. order ammonoidea zittel 1884 suborder phylloceratina arkell 1950 superfamily phylloceratoidea zittel 1884 family phylloceratidae zittel 1884 subfamily phylloceratinae zittel 1884 genus phylloceras suess 1866 type species. ammonites heterophyllus j. sowerby 1820, p. 119, plate 26, by monotypy. subgenus phylloceras (hypophylloceras) salfeld 1924 type species. phylloceras onoense stanton 1895, p. 74, by monotypy. north sea basin north africa 500 km exposed land shallow marine deep marine 1 2 3 present-day coastlines greenland fennoscandia russian platform tethys atlantic fig. 4 simplified early cenomanian (100.5–95.7 ma) palaeogeography of europe and adjacent regions. some partial present-day coastlines are given for orientation where appropriate. occurrences of schloenbachia varians are marked by stars – those reported herein in red (1: hold with hope. 2: geographical society ø. 3: kangerlussuaq basin), previously published sites in green. map modified from wilmsen & mosavinia (2010) and kennedy (2013). https://doi.org/10.34194/geusb.v57.8366 http://www.geusbulletin.org/ kennedy et al. 2024: geus bulletin 57. 8366. https://doi.org/10.34194/geusb.v57.8366 7 of 29 geusbulletin.org phylloceras (hypophylloceras) lombardensis (joly 2000) figs 5a–d ?1954 phylloceras cf. velledae (michelin); donovan, p. 5. 2000 hyporbulites lombardensis joly, p. 163, plate 39, fig. 11; text-fig. 387. 2009 hyporbulites lombardensis joly 2000; klein et al. pp. 90, 92. type. the holotype, by original designation, is in the b. joly collection (presumably in private hands; no repository given) and is from the condensed lower and middle cenomanian banc des lombards, cassis, bouches-durhône, france. material. mguh 34489 and mguh 34467 from casp locality w4225, ‘windy valley’, kangerlussuaq basin. description. mguh 34489 (figs 5a, b) is a 120° sector of phragmocone retaining traces of recrystallised shell, with a maximum preserved whorl height of 39 mm. coiling is very involute, the tiny umbilicus with a convex, outward-inclined umbilical wall and rounded umbilical shoulder. the whorl section is compressed subrectangular, with a whorl breadth to height ratio of 0.64, the greatest breadth just outside the umbilical shoulder. the flanks are flattened and subparallel, converging ventrally. the ventrolateral shoulders are broadly rounded, the venter feebly convex. delicate riblets arise as mere striae on the umbilical wall, and strengthen across the umbilical shoulder, where they are grouped in bundles. the riblets are very feebly convex across the umbilical shoulders, prorsiradiate and near straight on the flanks, flexing very feebly forwards to cross the venter in the feeblest of convexities. mguh 34467 (figs 5c, d) is a much larger internal mould fragment of the outermost flank, ventrolateral shoulder, and venter of a phragmocone. the riblets are very feebly prorsiradiate on the flanks and ventrolateral shoulder, and near-transverse over the venter. discussion. these specimens are referred to p. (h.) lombardensis on the basis of whorl section and ornament, with the riblets tending to arise in bunches from the umbilical shoulder. occurrence. condensed lower or middle cenomanian of cassis, bouches-du-rhône, france, and east greenland. suborder lytoceratina hyatt 1889 superfamily tetragonitoidea hyatt 1900 family gaudryceratidae spath 1927 genus and subgenus gaudryceras de grossouvre 1894 type species. ammonites mitis hauer 1866, p. 305, plate 2, figs 3, 4, by the subsequent designation of boule et al. 1906, p. 183(11). gaudryceras (gaudryceras) cassisianum (d’orbigny 1850) figs 6a–f 1850 ammonites cassisianus d’orbigny 1850, p. 146. 1940 gaudryceras cassisianum d’orbigny sp.; fabre, p. 216, plate 5, figs 8, 9. ?1954 gaudryceras sp. indet. donovan, p. 7, plate 2, fig. 3. 1955 ammonites (gaudryceras) cassisianus d’orbigny; sornay, fiche 25, figs 1–3. 1978 anagaudryceras cf. cassisianum (d’orbigny 1850); cooper, p. 65, fig. 8, non figs 9, 10, 11b-d. fig. 5 phylloceras (hypophylloceras) lombardense (joly 2000). a, b: mguh 34489. c, d: mguh 34452; both from casp locality w4225, ‘windy valley’, kangerlussuaq. a, b are ×2; c, d are ×1. https://doi.org/10.34194/geusb.v57.8366 http://www.geusbulletin.org/ kennedy et al. 2024: geus bulletin 57. 8366. https://doi.org/10.34194/geusb.v57.8366 8 of 29 geusbulletin.org ?1980 anagaudryceras cf. cassisianum (d’orbigny 1850); marcinowski, p. 243, plate 1, fig. 13. 1987 gaudryceras cenomanense thomel, p. 3, plate 1, fig. 6. 1992 gaudryceras cenomanense thomel; thomel, plate 7, figs 6–9; text-fig. 30h. 1994 gaudryceras (gaudryceras) cassisianum (d’orbigny 1850); kennedy, p. 218, plate 2, figs 8, 9; plate 3, figs 9–13, 15, 16, 18, 19, 24; plate 12, fig. 13. 2009 gaudryceras cassisianum (d’orbigny 1850); klein et al. pp. 172, 175 (with additional synonymy). types. the lectotype, by the subsequent designation of sornay (1955) was refigured by kennedy (1994, plate 3, figs 9, 10), as was the paralectotype (1994, plate 3, fig. 12). the specimens are both registered as 6122 in the d’orbigny collection, housed in the muséum national d’histoire naturelle, paris. material. two specimens, mguh 34449 from casp locality w4225, ‘windy valley’, kangerlussuaq basin and mguh 34483 from casp locality k7351, tværdal, geographical society ø. description. mguh 34449 (fig. 6a–d) is a 120° sector of phragmocone with extensive areas of recrystallised shell preserved. coiling is evolute, with a broad umbilicus of moderate depth. the whorl section is circular and slightly depressed. delicate lirae and striae arise on the umbilical wall, sweep back and strengthen across the umbilical shoulder, are prorsiradiate, and very feebly sinuous across the flanks and transverse over the venter. two strong collars are present on the fragment. they follow the course of the riblets and striae and strengthen progressively across flanks, ventrolateral shoulders, and venter. discussion. kennedy (1994) described and illustrated the types and a range of topotypes that differ in no significant respects from the present specimens. gaudryceras cenomanense thomel 1987 (p. 3, plate 1, fig. 6) is a clear synonym. occurrence. the types are from the condensed lower middle cenomanian banc des lombards of cassis, bouches-du-rhône, france. the species also occurs in the middle cenomanian of saint-lyons, alpes de hauteprovence, france, and east and north-east greenland. subgenus mesogaudryceras spath 1927 type species. ammonites leptonema sharpe 1855, p. 32, plate 14, fig. 3, by the original designation of spath (1927, p. 66). gaudryceras (mesogaudryceras) leptonema (sharpe 1855) figs 7a–o 1855 ammonites leptonema sharpe, p. 32, plate 14, fig. 3. 1953 mesogaudryceras cf. leptonema (sharpe); donovan, p. 120. fig. 6 gaudryceras (gaudryceras) cassisianum (d’orbigny 1850). a–d: mguh 34449, casp locality no. w4225, ‘windy valley’, kangerlussuaq. e, f: mguh 34483, casp locality no. k7351, tvaerdal, geographical society ø. all ×2. https://doi.org/10.34194/geusb.v57.8366 http://www.geusbulletin.org/ kennedy et al. 2024: geus bulletin 57. 8366. https://doi.org/10.34194/geusb.v57.8366 9 of 29 geusbulletin.org fig. 7 gaudryceras (mesogaudryceras) leptonema (sharpe 1855). a–d: mguh 34488. e: mguh 34466. f–h: mguh 34457. i–k: mguh 34487. l, m: mguh 34448; all casp locality no. w4225, ‘windy valley’, kangerlussuaq basin. n: mguh 34471. o: mguh 34472; both from casp locality no. w4342, leitch bjerg, geographical society ø. all ×2. https://doi.org/10.34194/geusb.v57.8366 http://www.geusbulletin.org/ kennedy et al. 2024: geus bulletin 57. 8366. https://doi.org/10.34194/geusb.v57.8366 10 of 29 geusbulletin.org fig. 8 titanoleioceras boreale gen. et sp. nov. a–c: paratype, mguh 34464. d–f: paratype, mguh 34455. g–i: paratype, mguh 34465. j: paratype, mguh 34451. k–n: paratype, mguh 34456. o–q: paratype, mguh 34450; all from casp locality no. w4225, ‘windy valley’, kangerlussuaq basin. all ×2. https://doi.org/10.34194/geusb.v57.8366 http://www.geusbulletin.org/ kennedy et al. 2024: geus bulletin 57. 8366. https://doi.org/10.34194/geusb.v57.8366 11 of 29 geusbulletin.org 1954 mesogaudryceras cf. leptonema (sharpe); donovan, p. 23. ?1954 lytoceras sp. nov. cf. vicinum h. douvillé; donovan, p. 6, plate 2, fig. 8. 1984 gaudryceras (mesogaudryceras) leptonema (sharpe 1855); wright & kennedy, p. 51, plate 2, fig. 8; text-figs 3a-m (with full synonymy). 1987 mesogaudryceras leptonema (sharpe 1856); thomel, p. 10, plate 1, figs 1, 5, 12; plate 2, fig. 7; plate 3, fig. 1; plate 4, fig. 7; text-figs 3, 4, 5. 1992 mesogaudyrceras leptonema (sharpe); thomel, plate 7, figs 4, 5; plate 10, figs 6, 7, 9–11; textfig. 30d, e. 1994 gaudryceras (mesogaudryceras) leptonema (sharpe 1855); kennedy, p. 218, plate 3, figs 4–8, 14, 23. 1997 mesogaudryceras leptonema (sharpe 1855); arkadiev & bogdanova, p. 110, plate 46, figs 1, 2. 2000 mesogaudryceras leptonema (sharpe 1855); arkadiev et al. p. 99, plate 10, figs 1, 2 2009 gaudryceras leptonema (sharpe 1855): klein et al. pp. 171, 183. type. the holotype, by monotypy, is no. 7762 in the collections of the british geological survey, keyworth, from the cenomanian lower chalk of ventnor, isle of wight, uk, the original of sharpe (1855, plate 14, fig. 3), refigured by wright & kennedy (1984, plate 2, figs 8a-c). material. eight specimens, mguh 34448, mguh 34457, mguh 34466, mguh 34487, mguh 34488, gm 2024.2, gm 2024.5 and gm 2024.6, from casp locality w4225, ‘windy valley’, kangerlussuaq, and three specimens, mguh 34471, mguh 34471 and gm 2024.9, from casp locality w4342, leitch bjerg, geographical society ø. description. phragmocones are up to 30.4 mm in diameter. coiling is evolute, the broad umbilicus comprising up to 33% of the diameter, shallow, with a convex umbilical wall and shoulder. the whorl section is circular to slightly compressed (whorl breadth to height ratios down to 0.92). internal moulds are smooth (mguh 34487; fig. 7i–k). most specimens retain a recrystallised shell, the surface of which is ornamented by wiry primary ribs that are much narrower than the interspaces. the ribs arise at the umbilical seam and strengthen across the umbilical wall (where they are markedly prorsiradiate). the ribs are prorsiradiate, strengthening progressively across the flanks, ventrolateral shoulders and venter, sinuous, feebly convex on ventrolateral shoulder and innermost flank, convex on the outer flank, sweeping forwards across the ventrolateral shoulder, and crossing the venter in a marked convexity. fragments of body chamber associated with mguh 34457 (fig. 7f–h), together with gm 2024.2 (not figured), show this style of ornament extending to estimated whorl heights of up to 20 mm. in some specimens, occasional ribs are strengthened and form feeble flares (mguh 34488: fig. 7a–d). the septal lobe is very large (fig. 7f, i). discussion. the present specimens differ in no significant respects from the holotype, and material from southeast france described by thomel (1987, 1992), wright & kennedy (1984), and kennedy (1994). mesogaudryceras rarecostatus balan 1979 (p. 40, fig. 2; marcinowski 1980, p.  244, plate 1, figs 9, 10b, 11, 12), originally described from the cenomanian of ukraine (crimea), is said to have more widely spaced ribs than g. (m.) leptonema, of which it may well be a synonym. occurrence. lower or middle cenomanian of southern england. lower and middle cenomanian of southeast france, the bavarian alps, middle cenomanian of ukraine (crimea), kopet dag, turkmenistan and east and northeast greenland. 5 mm e l u2 u1 i u3 = s fig. 9 suture of a juvenile titanoleioceras boreale gen. et sp. nov., mguh 34455, casp locality no. w4225, ‘windy valley’, kangerlussuaq basin. e: external lobe. a: adventive lobe. u1, u2, u3: umbilical lobes. i: internal lobe. l: lateral lobe. https://doi.org/10.34194/geusb.v57.8366 http://www.geusbulletin.org/ kennedy et al. 2024: geus bulletin 57. 8366. https://doi.org/10.34194/geusb.v57.8366 12 of 29 geusbulletin.org family tetragonitidae hyatt 1900 subfamily tetragonitinae hyatt 1900 genus titanoleioceras gen. nov. derivation of name. titan (greek), a giant, son of uranus and gaia; leios (greek), smooth; keros (greek), horn. hence giant smooth horn. type species. titanoleioceras boreale sp. nov. diagnosis. giant tetragonitinae with phragmocones up to 230 mm in diameter. whorl section as wide as high or depressed, trapezoidal. smooth but for growth lines and striae during early growth stages. fig. 10 titanoleioceras boreale gen. et sp. nov., holotype, mguh 34458, casp locality no. w4225, ‘windy valley’, kangerlussuaq basin; ×1. https://doi.org/10.34194/geusb.v57.8366 http://www.geusbulletin.org/ kennedy et al. 2024: geus bulletin 57. 8366. https://doi.org/10.34194/geusb.v57.8366 13 of 29 geusbulletin.org feeble, very widely separated constrictions appear in middle growth and are prominent on the adult body chamber; straight and prorsiradiate on the flanks, flexed back on ventrolateral shoulders and feebly concave over venter. striations develop on the adult body chamber. suture (fig. 9) with irregularly trifid saddles and bifid lobes and small auxiliary saddles; septal lobe massive. suture becomes highly subdivided at maturity. discussion. titanoleioceras is interpreted as a hypermorphic giant derivative of tetragonites, reaching a similar size to the largest species of pseudophyllites kossmat 1895. it differs from that genus in the persistence of constrictions fig. 11 titanoleioceras boreale gen. et sp. nov., holotype, mguh 34458, casp locality no. w4225, ‘windy valley’, kangerlussuaq basin; ×1. https://doi.org/10.34194/geusb.v57.8366 http://www.geusbulletin.org/ kennedy et al. 2024: geus bulletin 57. 8366. https://doi.org/10.34194/geusb.v57.8366 14 of 29 geusbulletin.org fig. 12 titanoleioceras boreale gen. et sp. nov., holotype, mguh 34458, casp locality no. w4225, ‘windy valley’, kangerlussuaq; ×1. to maturity: they are absent in pseudophyllites. takahashia matsumoto 1984 (type species t. eurekae matsumoto 1984, p. 33, text-fig. 1) is known from the holotype only, an individual 78 mm in diameter. it lacks constrictions and has well-developed crowded ribs that are markedly convex on the flanks. occurrence. lower cenomanian, east greenland. titanoleioceras boreale gen et sp. nov. figs 8a–q, 9–16 derivation of name. from boreas (greek), north. https://doi.org/10.34194/geusb.v57.8366 http://www.geusbulletin.org/ kennedy et al. 2024: geus bulletin 57. 8366. https://doi.org/10.34194/geusb.v57.8366 15 of 29 geusbulletin.org fig. 13 titanoleioceras boreale gen. et sp. nov., paratype, mguh 34459, casp locality no. w4225, ‘windy valley’, kangerlussuaq; ×1. types. the holotype is mguh 34458 (figs 10–12). there are nine paratypes, mguh 34450, mguh 34451, mguh 34454, mguh 34455, mguh 34456, mguh 34459, and mguh 34463–34465. type locality and horizon. all specimens are from the casp locality w4225, ‘windy valley’, kangerlussuaq basin (coordinates: 68.6383, –30.9297 wgs84 datum), from the sorgenfri formation. several specimens of titanoleioceras boreale occur below and above a horizon that yielded schloenbachia varians, which confirms their early cenomanian age (fig. 3). however, most specimens were collected from scree at the base of the section. diagnosis. with the characters of the genus. dimensions. measurements on four paratypes (mguh 34465, mguh 34455, mguh 34456, and mguh 34454) https://doi.org/10.34194/geusb.v57.8366 http://www.geusbulletin.org/ kennedy et al. 2024: geus bulletin 57. 8366. https://doi.org/10.34194/geusb.v57.8366 16 of 29 geusbulletin.org fig. 14 titanoleioceras boreale gen. et sp. nov., paratype, mguh 34459, casp locality no. w4225, ‘windy valley’, kangerlussuaq basin; ×1. https://doi.org/10.34194/geusb.v57.8366 http://www.geusbulletin.org/ kennedy et al. 2024: geus bulletin 57. 8366. https://doi.org/10.34194/geusb.v57.8366 17 of 29 geusbulletin.org fig. 15 titanoleioceras boreale gen. et sp. nov., paratype, mguh 34454, casp locality no. w4225, ‘windy valley’, kangerlussuaq basin; ×1. https://doi.org/10.34194/geusb.v57.8366 http://www.geusbulletin.org/ kennedy et al. 2024: geus bulletin 57. 8366. https://doi.org/10.34194/geusb.v57.8366 18 of 29 geusbulletin.org fig. 16 titanoleioceras boreale gen. et sp. nov., paratype, mguh 34454, casp locality no. w4225, ‘windy valley’, kangerlussuaq; ×1. https://doi.org/10.34194/geusb.v57.8366 http://www.geusbulletin.org/ kennedy et al. 2024: geus bulletin 57. 8366. https://doi.org/10.34194/geusb.v57.8366 19 of 29 geusbulletin.org fig. 17 parapuzosia (austiniceras) austeni (sharpe 1855). a–c: mguh 34460. d, e: mguh 34461; both from casp locality no. w4225, ‘windy valley’, kangerlussuaq basin; ×1. https://doi.org/10.34194/geusb.v57.8366 http://www.geusbulletin.org/ kennedy et al. 2024: geus bulletin 57. 8366. https://doi.org/10.34194/geusb.v57.8366 20 of 29 geusbulletin.org fig. 18 parapuzosia (austiniceras) austeni (sharpe 1855). a: mguh 34460. b–d: mguh 34462; both from casp locality no. w4225, ‘windy valley’, kangerlussuaq basin; ×1. https://doi.org/10.34194/geusb.v57.8366 http://www.geusbulletin.org/ kennedy et al. 2024: geus bulletin 57. 8366. https://doi.org/10.34194/geusb.v57.8366 21 of 29 geusbulletin.org are shown here. numbers in parentheses are dimensions as a percentage of diameter. mguh no. d wb wh wb:wh u mguh 34465 (-) 11.2 (-) 9.7 (-) 1.15 (-) mguh 34455 (-) 20.5 (-) 20.5 (-) 1.0 (-) mguh 34456 27.5 (100) 17.6 (64.0) 15.4 (56.0) 1.14 4.5 (16.4) mguh 34454 225.0 (100) 133.0 (59.1) 130.3 (57.8) 1.02 29.5 (13.1) d, wb, wh, u in mm.-: not measurable description. a series of juveniles have whorl heights of up to 25 mm. coiling is very involute, the umbilicus deep, with a flat, outward-inclined wall, and relatively narrowly rounded umbilical shoulder. the whorl section is as broad as high to slightly depressed (whorl breadth to height ratios of up to 1.15), with flattened, convergent flanks, broadly rounded ventrolateral shoulders, and a broad, feebly convex venter (figs 8a–q). the surface of the internal mould is smooth. the penultimate whorl of the holotype retains traces of replaced shell, with an ornament of prorsiradiate growth lines and striae only on the flanks. there is a single feeble constriction on the 180° sector of whorl, prorsiradiate on the flanks, flexing back and feebly convex across the ventrolateral shoulders, and very feebly concave over the venter. paratype mguh 34454 (figs 15, 16) is a phragmocone 225 mm in diameter. the penultimate whorl is smooth. the outer whorl of the specimen bears a subdued ornament of low, effaced strigations, most obvious on the venter, and low narrow undulations on the outer flanks and venter. paratype mguh 34459 (figs 13, 14) is a half whorl 240 mm fig. 19 parapuzosia (austiniceras) austeni (sharpe 1855). a, b: mguh 34452, casp locality no. w4225, ‘windy valley’, kangerlussuaq basin; ×1. https://doi.org/10.34194/geusb.v57.8366 http://www.geusbulletin.org/ kennedy et al. 2024: geus bulletin 57. 8366. https://doi.org/10.34194/geusb.v57.8366 22 of 29 geusbulletin.org in diameter, with a single feeble constriction towards the adapertural end. the outer whorl of the holotype (figs 10–12) is a 120° sector from the adapical end of the adult body chamber. two prominent constrictions are present, one towards the adapical end, the other at the adapertural end. they are very feebly concave on the umbilical wall, prorsiradiate on the flanks, across which they broaden and deepen, flex back, and are feebly convex on the ventrolateral shoulder and cross the venter in a very feeble convexity. growth lines and shallow grooves parallel the constrictions, and feeble strigations are present, most conspicuous on the ventrolateral shoulders and venter where the shell is preserved but effaced on the internal mould. the suture is described in the diagnosis for the genus. discussion. differences from species of pseudophyllites kossmat 1895 and takahashia matsumoto 1984, are discussed with the genus. occurrence. as for genus. suborder ammonitina hyatt 1889 superfamily desmoceratoidea zittel 1895 family desmoceratidae zittel 1895 subfamily puzosiinae spath 1922 genus parapuzosia nowak 1913 type species. sonneratia daubréei grossouvre 1894, p. 154, plate 28; by original designation. subgenus parapuzosia (austiniceras) spath 1922 type species. ammonites austeni sharpe 1855, p. 28, plate 12, fig. 1; by original designation by spath (1922, p. 127). parapuzosia (austiniceras) austeni sharpe 1855 figs 17a–e, 18a–d, 19a, b, 20 1855 ammonites austeni sharpe, p. 28, plate 12, fig.  1, non 2 [= puzosia (anapuzosia) dibleyi (spath 1922)]. 1984 parapuzosia (austiniceras) austeni (sharpe 1855); wright & kennedy, p.  60, plate  5,figs  3, 6; text-fig. 5 (with full synonymy). 10 mm l u2 fig. 20 partial external suture of parapuzosia (austiniceras) austeni (sharpe 1855), mguh 34452, casp locality no. w4225, ‘windy valley’, kangerlussuaq basin. l: lateral lobe. u2: umbilical lobe. https://doi.org/10.34194/geusb.v57.8366 http://www.geusbulletin.org/ kennedy et al. 2024: geus bulletin 57. 8366. https://doi.org/10.34194/geusb.v57.8366 23 of 29 geusbulletin.org fig. 21 schloenbachia varians (j. sowerby 1817). a, b: mguh 34453, casp locality no. w4225, ‘windy valley’, kangerlussuaq basin. c–e: mguh 34485. f, g: mguh 34484; both from casp locality no. w3099, tvaerdal, geographical society ø. h, i: mguh 34482, casp locality no. k7350, tvaerdal, geographical society ø. j, k: mguh 34490, casp locality no. w3099, tvaerdal, geographical society ø. all ×1. https://doi.org/10.34194/geusb.v57.8366 http://www.geusbulletin.org/ kennedy et al. 2024: geus bulletin 57. 8366. https://doi.org/10.34194/geusb.v57.8366 24 of 29 geusbulletin.org fig. 22 schloenbachia varians (j. sowerby 1817). a: mguh 34479. b: mguh 34475; both casp locality no. w4390, lysdal, geographical society ø. c: mguh 34473, casp locality no. w4389, lysdal, geographical society ø. d: mguh 34477, casp locality no. w4390, lysdal, geographical society ø. e: mguh 34469, casp locality no. w4342, leitch bjerg, geographical society ø. f: mguh 34480, casp locality no. w4390, lysdal, geographical society ø. g: mguh 34468, casp locality no. w4342, leitch bjerg, geographical society ø. h: mguh 34476; i: mguh 34474; both casp locality no. w4390, lysdal, geographical society ø. j: mguh 34470, casp locality no. w4342, leitch bjerg, geographical society ø. k: mguh 34478, casp locality no. w4390, lysdal, geographical society ø. l: mguh 34481 (silicone squeeze taken from external mould), casp locality no. k7350, tværdal, geographical society ø. m: mguh 34486, casp locality no. k7359, lygnaelv, hold with hope. all ×1. https://doi.org/10.34194/geusb.v57.8366 http://www.geusbulletin.org/ kennedy et al. 2024: geus bulletin 57. 8366. https://doi.org/10.34194/geusb.v57.8366 25 of 29 geusbulletin.org 1987 parapuzosia (austiniceras) austeni (sharpe); wright & kennedy, p. 144, text-fig. 7.1. non 1992 parapuzosia (austiniceras) austeni (sharpe); thomel, p. 209, plate 82 [= puzosia (anapuzosia) sp.]. 1998 parapuzosia (austiniceras) austeni (sharpe 1855); kaplan et al. p. 78, plate 1, figs 4, 7, 9; plate 2; plate 4. types. the lectotype, designated by spath (1922, p. 127) is c3382 in the collections of the natural history museum, london, from the lower chalk of guildford, surrey; it is the original of sharpe 1855, plate 12, figs1a, b. paralectotypes have not been traced. material. seven specimens, mguh 34452, mguh 34460, mguh 34461, mguh 34462, gm 2024.3, gm 2024.4, and gm 2024.7 from casp locality w4225, ‘windy valley’, kangerlussuaq. description. mguh 34462 (fig. 18b–d) may be a microconch phragmocone with an original estimated diameter of 100–110 mm, retaining traces of recrystallised shell. coiling is evolute, with 48% of the previous whorl covered, the umbilicus broad, shallow, with a low, feebly convex wall, and narrowly rounded umbilical shoulder. the whorl section is compressed, with a whorl breadth to height ratio of 0.7. the greatest breadth is below mid-flank. the inner flanks are flattened, the middle and outer flanks feebly convex, converging to broadly arched ventrolateral shoulders and venter. there appear to be as many as eight constrictions on the inner flanks of the penultimate whorl. these are narrow, straight, and feebly prorsiradiate, with no other ornament on the surface of the intervening internal mould. on the outer whorl of the specimen, there are three constrictions per half whorl, deeply incised into the umbilical wall, strong and narrow on the flanks, straight and prorsiradiate on the inner flank, flexing back at mid-flank, feebly concave on the outer flank, sweeping forward across the ventrolateral shoulders, and crossing the venter in an obtuse linguoid peak. the adapical edge of the constrictions is strengthened into a feeble collar-rib. there are up to 16 ribs between successive constrictions. these parallel the constrictions, arising as mere striae on the inner flank, but strengthening on outer flanks, ventrolateral shoulders and venter, which bear an ornament of crowded, even, rounded ribs that are concave on outer flanks and ventrolateral shoulder and cross the venter in a rounded linguoid peak. at the adapertural end of the specimen, the last constriction and succeeding ribs modify their course, projecting strongly forward on the inner flank and back at mid-flank, to define the beginnings of lateral lappets (fig. 18b). traces of a further 240° of body chamber are preserved (fig. 18c). mguh 34461 (fig. 17d, e) is a well-preserved phragmocone fragment with recrystallised shell preserved; the maximum preserved whorl height is 42 mm. the whorl breadth to height ratio is 0.74, the inner flanks are feebly convex and subparallel, the middle and outer flanks converging to the broadly rounded venter. parts of two flared ribs are preserved, marking the adapical edge of constrictions on the internal mould. these collars are well-developed on the outer flanks, ventrolateral shoulders and venter, which they cross in an obtuse linguoid peak. ornament is reduced to growth lines, striae, and riblets between the collars. there are fine, even riblets on outer flanks, ventrolateral shoulders and venter on the adapertural end of the fragment, following a narrow, smooth zone associated with the site of the adapertural of the two constrictions. gm 2024.3 (not figured) is a corroded phragmocone fragment with a maximum preserved whorl height of around 70 mm; it overlaps in size with mguh 34452 (figs 19a, b), a wholly septate fragment with traces of recrystallised shell preserved and the following approximate dimensions: d = 160(100); wb = 49.8(31.1); wh = 71.4(44.6); wb:wh = 0.7; u = 48.7(30.4). the umbilicus is shallow, with a flattened, outward-inclined wall and very narrowly rounded umbilical shoulder. the flanks are flattened and subparallel, the ventrolateral shoulders and venter broadly rounded. there are two constrictions, and possibly a third at the adapical end. these are conspicuous on the internal mould, straight and feebly prorsiradiate on the inner flank, flexing back and concave at midflank, markedly concave on the outer flank, projecting forwards on the ventrolateral shoulders and effacing markedly on the venter, where they form an obtuse chevron. ornament is very subdued on the internal mould. where recrystallised shell is preserved, there is a strong adapical collar rib preserved, associated with the adapertural constriction. the shell surface between constrictions or collars is ornamented by more than 30 crowded, even ribs. these are straight, prorsiradiate, and weak on the inner flanks, flexing back and convex at mid-flank, strengthening and concave on the outer flank, and broadly convex and at their maximum strength over the venter. gm 2024.4 and gm 2024.7 (not figured) are small fragments of similar-sized individuals. mguh 34460 (figs 17a–c, 18a) is a further phragmocone fragment, with a maximum preserved whorl height of 110 mm. the whorl breadth to height ratio is 0.67. the umbilicus is shallow, the umbilical wall flattened and inclined outwards. the umbilical shoulder is very narrowly rounded, and sharply defined. the inner to middle flank region is feebly convex, the outer flanks convergent, the ventrolateral shoulders broadly https://doi.org/10.34194/geusb.v57.8366 http://www.geusbulletin.org/ kennedy et al. 2024: geus bulletin 57. 8366. https://doi.org/10.34194/geusb.v57.8366 26 of 29 geusbulletin.org rounded, the venter very feebly convex (fig. 17b). the surface of the recrystallised shell is ornamented by crowded even ribs. these arise on the umbilical shoulder. they are straight and feebly prorsiradiate on the inner flank, flexed back and feebly convex at mid-flank, flexing back, and concave on the outer flanks. additional short ribs intercalate on the outermost flanks, and all ribs strengthen and sweep forwards on the ventrolateral shoulders and cross the venter in a very broad convexity. the suture (figs 19a, 20) is deeply and intricately subdivided, with a strongly retracted suspensive lobe. discussion. see wright & kennedy (1984, p. 60). occurrence. lower cenomanian to middle turonian. the geographic distribution extends from southern england to germany, france, the czech republic, ukraine (crimea), east greenland, and perhaps kwazulu-natal (south africa). superfamily hoplitoidea h. douvillé 1890 family schloenbachiidae parona & bonarelli 1897 genus schloenbachia neumayr 1875 type species. ammonites varians j. sowerby 1817, p. 169, plate 176, by the subsequent designation of h. douvillé 1890, p. 200. schloenbachia varians (j. sowerby 1817) figs 21a–k, 22a–m 1817 ammonites varians j. sowerby, p. 169 (pars), plate 176 uppermost figure, three figures in row below, right-hand figure in row below. 1953 schloenbachia subvarians spath; donovan, p. 119. 1953 schloenbachia subtuberculata (sharpe); donovan, p. 119. 1953 schloenbachia cf. subplana (mantell); donovan, p. 120. 1954 schloenbachia aff. subplana (mantell); donovan, p. 9. 1954 schloenbachia subvarians spath; donovan, p. 10, plate 1, figs 1, 2, 6, 7; text-fig. 2. 1954 schloenbachia cf. varians (j. sowerby); donovan, p. 12, plate 2, fig. 2. 1954 schloenbachia subvarians var. tetrammata (j. de c. sowerby); donovan, p. 12, plate 2, fig. 1. 1954 schloenbachia sp.; donovan, p. 12, plate 1, fig. 5. 1998 schloenbachia ventriosa stieler; kelly et  al. p. 1005. 2015 schloenbachia varians (j. sowerby 1817); kennedy in wright & kennedy, p. 419, plate 125, figs 2, 5; plates 126–135; plate 136, fig. 1; text-figs 161–163; 164b, c;165–167; 169–173 (with full synonymy). type. lectotype, by the subsequent designation of spath (1938, p. 544), is no 43962b in the collections of the natural history museum, london. the original of sowerby (1817, plate 176, top figure), refigured by kennedy in wright & kennedy (2015, plate 125, fig. 2). it is from the lower chalk at an unknown locality in southern england. material. mguh 34484, mguh 34485, mguh 34490, and gm 2024.10–2024.14 from casp locality w3099, and mguh 34481 and mguh 34482 from casp locality k7350, both tværdal, geographical society ø. mguh 34453 and gm 2024.8 from casp locality w4225, ‘windy valley’, kangerlussuaq basin. mguh 34468–34470 from casp locality w4342, leitch bjerg, geographical society ø. mguh 34473 from casp locality w4389, and mguh 34474–34480 from casp locality w4390, both lysdal, geographical society ø. mguh 34486 from casp locality k7285, hold with hope. description. the material is highly variable, as with populations from elsewhere. the formae recognised are, from strongly to weakly ornamented, ventriosa, varians, subtuberculata, intermedia, and subplana. compressed individuals, such as mguh 34484 (fig. 21f, g) and mguh 34453 (fig. 21a, b), have ribs that arise either singly or in pairs at the umbilical shoulder, with a very delicate umbilical tubercle and a slightly stronger inner lateral one, at which point the ribs commonly bifurcate, so that there are more ribs at the ventrolateral shoulder than at the umbilical shoulder. the ribs are straight and prorsiradiate on the innermost flank, but flex back and are feebly sinuous, feebly convex at mid-flank, and feebly concave on the outer flank. all ribs terminate in small ventral clavi on either side of the siphonal keel. these specimens correspond to the forma subvarians, although the presence of two tiny tubercles separates them from sharpe’s specimen (1853, plate 8, fig. 7), in which respect they are transitional to forma subtuberculata. this form is well-represented by mguh 34485 (fig. 21c–e), a specimen 66 mm in diameter, with stronger umbilical and inner lateral tuberculation, the ribs showing a tendency to loop between lateral tubercles and ventral clavi. other specimens of subvarians-subtuberculata type are gm 2024.8, gm 2024.10, and gm 2024.14 (not figured). gm 2024.11 (not figured) is an external mould of a crushed individual 80 mm in diameter with stronger ornaments and thus transitional between subtuberculata and the holotype of varians. mguh 34481 (fig. 22l) is a further very crushed macroconch of this type. the ornament is very well-preserved. on the inner whorls there are 11–12 small umbilical bullae, each linked by a low rib to large subspinose inner lateral bullae. the outer whorl https://doi.org/10.34194/geusb.v57.8366 http://www.geusbulletin.org/ kennedy et al. 2024: geus bulletin 57. 8366. https://doi.org/10.34194/geusb.v57.8366 27 of 29 geusbulletin.org is preserved to a height of 45 mm. coarser ribbed and tuberculate still are mguh 34482 and mguh 34490. the former (fig. 21h, i) may be an adult microconch. the penultimate whorl bears weak umbilical and very strong, coarse lateral tubercles, as in passage forms between formae varians and ventriosa. the fragmentary body chamber, preserved to a whorl height of 36 mm, appears to show eccentric coiling, the umbilical wall is flat and outward-inclined. small bullae perch on the umbilical shoulder and give rise to one or a pair of prorsiradiate ribs. lateral tubercles have effaced. mguh 34490 (fig. 21j, k) is a half whorl of body chamber with a maximum preserved diameter of 120 mm. coarse bullae perch on the umbilical shoulder and give rise to coarse prorsiradiate ribs. these link to strong inner lateral bullae that give rise to one or two ribs that sweep forward and link to coarse ventrolateral clavi. these decline towards the adapical end of the specimen and are replaced by narrow, sharp, strongly prorsiradiate ribs that efface before reaching the strong siphonal keel (fig. 21k). this specimen appears to be a near-complete adult macroconch of a variant between formae varians and ventriosa. discussion. specimens of schloenbachia dominate the cenomanian assemblages from east and north-east greenland, as they do throughout the eurasian part of the boreal realm. the present material comprises typical variants of early cenomanian varians following the interpretation of kennedy in wright & kennedy (2015, p. 408 et seq.). thus, three successive species are recognised: schloenbachia varians (j. sowerby 1817), which is restricted to the lower cenomanian; schloenbachia coupei (brongniart 1822) of the lower middle cenomanian acanthoceras rhotomagense zone, and schloenbachia lymensis spath 1926, which first appears in the upper middle cenomanian acanthoceras rhotomagense zone and has its acme in the lower upper cenomanian calycoceras (proeucalycoceras) guerangeri zone. differences between the species are set out in their diagnoses: schloenbachia varians: “adults range from 41 to at least 185 mm in adult diameter. the most strongly tuberculate individuals have lateral and ventrolateral tubercles only on most or all of the phragmocone; umbilical bullae are present on the adult body chamber. compressed individuals in which ribs dominate over tubercles have umbilical and inner lateral tubercles, the latter linked to approximately twice as many ventrolateral clavi by a single rib; zigzag and looped ribbing is uncommon. individuals with intermediate ornament may have outer flank ribs strengthened into weak or incipient bullae. suture deeply incised, with bifid e/a and a/u2; a trifid.” (kennedy in wright & kennedy 2015, p. 422). schloenbachia coupei (brongniart 1822): “adults range from 36 to 130 mm in diameter. all but the most feebly ornamented variants have umbilical, lateral and ventrolateral tubercles on the phragmocone and part or all of the body chamber, and a fourth row of outer lateral bullae may develop in near-adult and adult individuals. suture with moderately incised bifid e/a and a/u2; a trifid.” (kennedy in wright & kennedy 2015, p. 436). schloenbachia lymensis spath 1926: “a small species; adults range from 27 to 66 mm in diameter. strongly ornamented variants with umbilical and lateral bullae linked by a narrow bar-like rib on phragmocone, sometimes extending onto adapical end of adult body chamber, lateral bullae efface on body chamber, as do umbilical and ventrolateral clavi close to adult aperture. more compressed variants have straight, flexuous, falcoid or lautiform ribs and constrictions on the adult body chamber. the weakest ornamented variants lack lateral bullae on the phragmocone, the body chamber ornament reduced to delicate riblets and lirae, tiny ventrolatral clavi and strong constrictions. suture […] only moderately incised, with broad bifid e/a and a/u2; a narrow and trifid.” (kennedy in wright & kennedy 2015, p. 451). schloenbachia species show wide intraspecific variation (e.g. wilmsen & mosavinia 2010; kennedy 2013), and previous authors introduced a host of specific and varietal names. kaplan et al. (1998, p. 106) introduced the term forma as a non-linnean term for morphological variants, and this approach, as developed by kennedy (2013) and kennedy in wright & kennedy (2015) is followed here. occurrence. lower cenomanian (100.5–95.7 ma), east and north-east greenland, uk, northern ireland, germany, switzerland, poland, the czech republic, france (as far south as the alpes-maritimes), russia, ukraine, turkmenistan, iran (fig. 4). the record from tebessa in algeria (mendir et al. 2019, p. 239, plate 1, fig. 3) appears to be an acompsoceras hyatt 1903, some species of which develop a siphonal ridge (wright & kennedy 1987, text-fig. 43e). 5 simon kelly (16 october 1949 – 19 may 2023): an appreciation simon joined the cambridge arctic shelf programme (now known by its acronym, casp) as a consultant in 1984, conducting field-based research on svalbard, prior to working as a palaeontologist with the british antarctic survey (1988–1994). he returned to casp in 1994, conducting fieldwork and integrated biostratigraphic studies in azerbaijan, kazakhstan, arctic canada and, most notably, east and north-east greenland. simon’s principal research focused on the jurassic and cretaceous https://doi.org/10.34194/geusb.v57.8366 http://www.geusbulletin.org/ kennedy et al. 2024: geus bulletin 57. 8366. https://doi.org/10.34194/geusb.v57.8366 28 of 29 geusbulletin.org periods, specialising in molluscan biostratigraphy and palaeoecology. his research was especially influential in establishing stratigraphical schemes for east and northeast greenland. simon authored 70 scientific articles (see appendix in schneider & pointon 2023). simon was also curator of casp’s geological collections until 2020, overseeing their transfer into a purpose-built rock store that now bears his name. in 2000, simon was awarded the polar medal for his arctic and antarctic service. during the subsequent 19 years, he participated in a further 14 arctic field seasons for which he was awarded a prestigious second clasp to his polar medal as part of the uk government’s new year honours list 2023. acknowledgements kennedy thanks david sansom of the department of earth sciences, oxford, for drafting the suture line diagrams, and eliza howlett of the oxford university museum of natural history for technical support. caroline s. pickles (now ithaca energy, aberdeen, uk) and the late andrew g. whitham (casp, cambridge, uk) collected several of the specimens studied. the latter also compiled the log and took the photographs depicted in fig. 3. benoit vautravers (casp, cambridge, uk) drafted the map used in fig. 1 and provided support with data archiving and access. peter alsen (geus, copenhagen, denmark) and an anonymous reviewer provided comments that helped to improve the manuscript. arden r. bashforth (natural history museum of denmark, copenhagen, denmark) kindly assisted with collection numbers. additional information funding statement fieldwork in greenland supported by casp’s industry sponsors. author contributions w.j.k.: designed the study, performed taxonomy and wrote most of the text; took most of the photographs; systematic palaeontology is by w.j.k. only. s.r.a.k.: collected, assembled, and prepared the material; initiated the study, engaged in discussion of taxonomy, contributed to an early draft of the manuscript, and approved the figures. s.s.: compiled sample data and artwork, wrote the materials section, and contributed to the remainder of the text. competing interests none declared. additional files supplementary file s1 (an .xls file) contains sample metadata, including a full list of sample names, locality names, geographical coordinates, and geological setting. it is available at: https://doi.org/10.22008/ fk2/lm7lgg. this published work and the nomenclatural acts it contains have been registered in zoobank: http://zoobank.org/urn:lsid:zoobank. org:pub:4f83504a-7c33-418d-b812-04dc558f16e4 references arkadiev, v.v., atabekian, a.a., barabobshkin, e.y. & bogdanova, t.n. 2000: stratigraphy and ammonites of cretaceous deposits of south-west crimea. palaeontographica a255, 85–128. https://doi. org/10.1127/pala/255/2000/85 arkadiev, v.v. & bogdanova, t.n. 1997: atlas of the cretaceous fauna of the south-west crimea. saint petersburg: izdatelstvo sankt petersburgsko gornogo instituta. arkell, w.j. 1950: a classification of the jurassic ammonites. journal of paleontology 24, 354–364. balan, t.m. 1979: material for the systematics of the family gaudryceratidae spath. in: [biostratigraphical research on the mesozoic and cainozoic deposits of the southern edge of the russian platform]. 157 + 4 unnumbered pages, shtiintsa, kishinev, pp. 28–43 [in russian]. bjerager, m. et  al. 2020: cretaceous lithostratigraphy of north-east greenland. bulletin of the geological society of denmark 68, 37–93. https://doi.org/10.37570/bgsd-2020-68-04 boule, m., lemoine, p. & thévenin, a. 1906–1907: paléontologie de madagascar. iii céphalopodes crétacés des environs de diego-suarez. annales de paléontologie 1, 173–192 (1–20); 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tværdal (casp locality no. k7350, coordinates 72.9633, –23.0667; no. k7351, coordinates 73.0003, –22.9717; no. w3099, coordinates 72.958, –23.0452); lysdal (casp locality no. w4389, coordinates 72.8957, –22.807; no. w4390, coordinates 72.8965, –22.8122); leitch bjerg (casp locality no. w4342, coordinates 72.8555, –22.5023); and ‘windy valley’ (casp locality no. w4225; coordinates: 68.6383, –30.9297). fig. 2 albian to maastrichtian lithostratigraphy of the kangerlussuaq basin (data sources: larsen et al. 2005; nøhr-hansen 2012), geographical society ø (data sources: parsons et al. 2017; bjerger et al. 2020) and hold with hope (data sources: kelly et al. 1998; bjerger et al. 2020). fig. 3 photographs and schematic log of casp locality no. w4225, ‘windy valley’, kangerlussuaq basin. a: upper half of exposure, with prominent, yellowish fig. 4 simplified early cenomanian (100.5–95.7 ma) palaeogeography of europe and adjacent regions. some partial present-day coastlines are given for orientation where appropriate. occurrences of schloenbachia varians are marked by stars – those reported herein in red (1: hold with hope. 2: geographical fig. 5 phylloceras (hypophylloceras) lombardense (joly 2000). a, b: mguh 34489. c, d: mguh 34452; both from casp locality w4225, ‘windy valley’, kangerlussuaq. a, b are ×2; c, d are ×1. fig. 6 gaudryceras (gaudryceras) cassisianum (d’orbigny 1850). a–d: mguh 34449, casp locality no. w4225, ‘windy valley’, kangerlussuaq. e, f: mguh 34483, casp locality no. k7351, tvaerdal, geographical society ø. all ×2. fig. 7 gaudryceras (mesogaudryceras) leptonema (sharpe 1855). a–d: mguh 34488. e: mguh 34466. f–h: mguh 34457. i–k: mguh 34487. l, m: mguh 34448; all casp locality no. w4225, ‘windy valley’, kangerlussuaq basin. n: mguh 34471. o: mguh 34472; both from casp locality no. w4342, leitch bjerg, geographical society ø. all ×2. fig. 8 titanoleioceras boreale gen. et sp. nov. a–c: paratype, mguh 34464. d–f: paratype, mguh 34455. g–i: paratype, mguh 34465. j: paratype, mguh 34451. k–n: paratype, mguh 34456. o–q: paratype, mguh 34450; all from casp locality no. w4225, ‘windy valley’, kangerlussuaq basin. all ×2. fig. 9 suture of a juvenile titanoleioceras boreale gen. et sp. nov., mguh 34455, casp locality no. w4225, ‘windy valley’, kangerlussuaq basin. e: external lobe. a: adventive lobe. u1, u2, u3: umbilical lobes. i: internal lobe. l: lateral lobe. fig. 10 titanoleioceras boreale gen. et sp. nov., holotype, mguh 34458, casp locality no. w4225, ‘windy valley’, kangerlussuaq basin; ×1. fig. 11 titanoleioceras boreale gen. et sp. nov., holotype, mguh 34458, casp locality no. w4225, ‘windy valley’, kangerlussuaq basin; ×1. fig. 12 titanoleioceras boreale gen. et sp. nov., holotype, mguh 34458, casp locality no. w4225, ‘windy valley’, kangerlussuaq; ×1. fig. 13 titanoleioceras boreale gen. et sp. nov., paratype, mguh 34459, casp locality no. w4225, ‘windy valley’, kangerlussuaq; ×1. fig. 14 titanoleioceras boreale gen. et sp. nov., paratype, mguh 34459, casp locality no. w4225, ‘windy valley’, kangerlussuaq basin; ×1. fig. 15 titanoleioceras boreale gen. et sp. nov., paratype, mguh 34454, casp locality no. w4225, ‘windy valley’, kangerlussuaq basin; ×1. fig. 16 titanoleioceras boreale gen. et sp. nov., paratype, mguh 34454, casp locality no. w4225, ‘windy valley’, kangerlussuaq; ×1. fig. 17 parapuzosia (austiniceras) austeni (sharpe 1855). a–c: mguh 34460. d, e: mguh 34461; both from casp locality no. w4225, ‘windy valley’, kangerlussuaq basin; ×1. fig. 18 parapuzosia (austiniceras) austeni (sharpe 1855). a: mguh 34460. b–d: mguh 34462; both from casp locality no. w4225, ‘windy valley’, kangerlussuaq fig. 19 parapuzosia (austiniceras) austeni (sharpe 1855). a, b: mguh 34452, casp locality no. w4225, ‘windy valley’, kangerlussuaq basin; ×1. fig. 20 partial external suture of parapuzosia (austiniceras) austeni (sharpe 1855), mguh 34452, casp locality no. w4225, ‘windy valley’, kangerlussuaq basin. l: lateral lobe. u2: umbilical lobe. fig. 21 schloenbachia varians (j. sowerby 1817). a, b: mguh 34453, casp locality no. w4225, ‘windy valley’, kangerlussuaq basin. c–e: mguh 34485. f, g: mguh 34484; both from casp locality no. w3099, tvaerdal, geographical society ø. h, i: mguh 34482, casp locality no. k7350, tvaerdal, geographical fig. 22 schloenbachia varians (j. sowerby 1817). a: mguh 34479. b: mguh 34475; both casp locality no. w4390, lysdal, geographical society ø. c: mguh 34473, casp locality no. w4389, lysdal, geographical society ø. d: mguh 34477, casp locality no. w4390, lysdal, geographical society ø. e: mguh 34469, casp locality no. w4342, leitch bjerg, geographical society ø. f: mguh 34480, casp locality no. w4390, lysdal, geographical society ø. g: mguh 34468, casp locality no. w4342, leitch bjerg, geographical society ø. h: mguh 34476; i: mguh 34474; both casp locality no. w4390, lysdal, table table 1 occurrences of ammonite species in the eight localities studied marked by ‘×’. kasia k. śliwińska et al: geus bulletin 57. 8365. https://doi.org/10.34194/geusb.v57.8365 1 of 18 review article miocene vegetation and climate in the eastern north sea basin, onshore denmark, compared to the present kasia k. śliwińska*1 , thomas denk2 , karen dybkjær1 , julie margrethe fredborg3 , sofie lindström3,4 , stefan piasecki5,6 , erik skovbjerg rasmussen6 1department of geo-energy and storage, geological survey of denmark and greenland (geus), copenhagen, denmark; 2swedish museum of natural history, stockholm, sweden; 3department of geosciences and natural resource management, university of copenhagen, copenhagen, denmark; 4geological survey of denmark and greenland (geus), copenhagen, denmark; 5globe institute, university of copenhagen, copenhagen, denmark; 6department of geophysics and sedimentary basins, geological survey of denmark and greenland (geus), copenhagen, denmark abstract despite often being referred to as a ‘coolhouse climate’, the climate during the miocene (23.03– 5.33 ma) was overall humid, warm and temperate. it was paced by orbitally driven cooler periods (the oligocene–miocene transition and mi-events) overprinted by a climatic optimum. global cooling during the late miocene brought more arid conditions with climate seasonality, which varied across western eurasia. sedimentary archives from onshore denmark comprise shallow marine siliciclastic deposits and discrete brown coal layers. hence, they allow us to infer past climates and environments using both marine and terrestrial fossils. the backbone for miocene stratigraphy and palaeoclimate reconstruction in the eastern north sea basin (present-day denmark) is the sønder (sdr.) vium sediment core, which penetrates a shallow marine succession and spans an interval from c. 22 to 8 ma. here, we present an improved age model for the core. during the miocene, forested lowlands predominated in the eastern north sea basin. coastal areas included rich riparian landscapes and delta areas of lignite-forming swamp forest. compilations of existing proxy records (pollen, spores, leaves, plant fragments and the organic biomarkers alkenones and membrane lipids) collectively show that the climate here was warm and moist during the early and middle miocene, while the late miocene was characterised by climate cooling and modernisation of the vegetation. the interval preceding the miocene climatic optimum was already warm and moist, and the onset was not characterised by a significant increase in temperature and precipitation. instead, the palynoflora indicates homogeneous vegetation and only a weak signal of warming shown by a minor increase of, for example, sabaloid palms and mastixiaceae. *correspondence: kksl@geus.dk received: 29 dec 2023 revised: 29 april 2024 accepted: 29 april 2024 published: 25 sep 2024 keywords: miocene, palaeoclimate, terrestrial, marine, north sea abbreviations brgdgt: branched glycerol dialkyl glycerol tetraethers ca: coexistence approach cmt: coldest month temperature map: mean annual precipitation mat: mean annual air temperature mco: miocene climatic optimum mfs: maximum flooding surface mmct: middle miocene climate transition mrs: maximum regressive surface omt: oligocene–miocene transition sb: sequence boundary sst: sea-surface temperature wmt: warmest month temperature geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: marit-solveig seidenkrantz (aarhus university, denmark) reviewed by: torsten utescher (senckenberg society for nature research, germany) and one anonymous reviewer funding: see page 14 competing interests: see page 15 additional files: see page 15 1 introduction the miocene (23.03–5.33 ma) witnessed the final transition from the ‘coolhouse’ climate of the oligocene (with a permanent ice sheet covering antarctica) to the modern ‘icehouse’ world characterised by bipolar glaciation. the coolhouse climate commenced c. 34 ma and the transition from the preceding ‘greenhouse’ climate was driven by a decline in atmospheric co2 and changes in ocean gateways and current systems (stickley et al. 2004; pearson et al. 2009; cristini et al. 2012; anagnostou et al. 2016; hutchinson et al. 2019). during the oligocene (34–23.03 ma), atmospheric co2 may have decreased to values as low as 400 ppm (rae et al. 2021) and several orbitally paced glaciations occurred in antarctica (wade & pälike 2004; pälike et al. 2006). in the north atlantic – north sea – nordic seas region, some of these events show a strong correlation with episodes of cooling of surface waters and decreasing sea levels (śliwińska et al. 2010; śliwińska & heilmann-clausen 2011; clausen et al. 2012; śliwińska 2019). one of the major glaciations and cooling events of the coolhouse world took place during the oligocene–miocene transition (omt), where global cooling coincided with a drop in sea level (miller et  al. 2005) marked by a https://doi.org/10.34194/geusb.v57.8365 https://orcid.org/0000-0001-5488-8832 https://orcid.org/0000-0001-9535-1206 https://orcid.org/0000-0002-8420-3379 https://orcid.org/0009-0004-7938-2730 https://orcid.org/0000-0001-8278-1055 https://orcid.org/0000-0002-7846-859x https://orcid.org/0000-0001-8603-8429 mailto:kksl@geus.dk https://creativecommons.org/licenses/by/4.0/deed.ast kasia k. śliwińska et al: geus bulletin 57. 8365. https://doi.org/10.34194/geusb.v57.8365 2 of 18 geusbulletin.org transient, large amplitude expansion of antarctic ice sheets. the associated glacial maximum is expressed by a positive shift of c. 1‰ in stable oxygen-isotope values (δ18o) measured in benthic foraminifera. this shift was one of the largest, abrupt (<1 myr) increases in benthic δ18o observed for the cenozoic (lear et  al. 2004). the occurrence of orbitally paced transient glacial episodes continued during the miocene (e.g. miller et  al. 1991; miller & mountain 1996). however, in the eastern north sea basin (present-day denmark), glacioeustatic sealevel fluctuations were superimposed on tectonically driven sea-level changes, making it difficult to decouple these two factors (rasmussen 2004). although the miocene was significantly cooler than the previous epochs of the cenozoic, the climate was relatively warmer and wetter than today; hence, it is more appropriately referred to as coolhouse than icehouse (westerhold et  al. 2020). atmospheric co2 in the miocene and the following pliocene might have reached 200–400 ppm (the cenozoic co2 proxy integration project (cenco2pip) consortium 2023). the type of coolhouse climate known from c. 34 to 3.5 ma (westerhold et al. 2020) was interrupted by an interval of transient warmth, known as the miocene climatic optimum (mco) c. 16.9–14.7 ma (e.g. shackleton & kennett 1975; zachos et  al. 2001). the mco was the warmest period over the last 23 myr (zachos et  al. 2001) associated with a minimum ice volume in antarctica (holbourn et al. 2007). the warming was caused by elevated atmospheric co2, most possibly caused by intense volcanism primarily from the columbia river basalts eruptions (e.g. kasbohm & schoene 2018). the mco was the last time in earth’s history where atmospheric co2 was as high as it is today, and possibly higher at c. 500 ppm (the cenozoic co2 proxy integration project (cenco2pip) consortium 2023). with atmospheric co2 > 420 ppm and global mean temperatures c. 7–8°c warmer than present, the mco falls in the middle range of climate states predicted for representative concentration pathway (rcp) 4.5 and rcp 6.0 (collins et al. 2013; where 4.5 and 6.0 are the stabilised levels of radiative forcing expected by 2100 according to different emission scenarios). therefore, the mco is considered one of the most accurate reference intervals for a possible future climate. up to 3°c of the warming observed in the middle (goldner et al. 2014) and late miocene (knorr et al. 2011) can be explained by the differences in palaeo-oceanography (continent position, seaways and ocean current systems), topography and vegetation when compared with the modern state. the remaining temperature increase of the mco is attributed to elevated atmospheric co2 clearly showing its strong impact on the global climate. today, atmospheric co2 has already surpassed 400 ppm, and so mid-pliocene warming at 400 ppm co2 is considered an optimistic ‘best-case scenario’ for the near-future. in this scenario, earth’s anthropogenic global temperature anomaly by 2100 is limited to 1.5–2°c (schellnhuber et al. 2016) above pre-industrial times. with atmospheric co2 >450 ppm, elevated global temperatures and a more intense hydrological cycle, the mco is often considered the most recent climatic optimum in earth’s history that is comparable with today’s climate. thus, following a moderate co2 emission scenario, it may be more realistic that we reach a similar climatic condition as the mco by 2100. despite being more critical for our understanding of tipping points in the global climate system under co2 emission scenarios >450 ppm, the transition towards the mco is poorly studied, and most existing studies of proxy records focus on the maximum. in contrast to the well-studied antarctic ice sheet, the existence and potential extent of the greenland ice sheet at the mco is highly uncertain (thiede et al. 2010). in europe, the only long-term palaeotemperature record across the middle miocene transition north of 55°n is from denmark. the mco was followed by a period of a major growth of the east antarctic ice sheet and an associated cooling, termed the middle miocene climate transition (mmct; 14.7–13.8 ma; flower & kennett 1994). the mmct was associated with large changes in the global carbon cycle and in the terrestrial biosphere, including aridification of midlatitude continental regions (e.g. flower & kennett 1994). increased stability of the antarctic ice sheet after 14.8 ma represents a crucial step in the establishment of the near-modern climate system. sea-surface temperature (sst) reached near-modern values 7–5.4 ma (herbert et  al. 2016). the late miocene cooling led to the development of c4 grasslands (cerling et  al. 1993). these large-scale climate-driven shifts in vegetation and landscape led to major turnovers in the terrestrial fauna favouring browsers feeding on grasses and shrubs (e.g. badgley et  al. 2008). changes in terrestrial animal and plant communities together with drier conditions across large areas of the continents gave rise to a world that is similar to that we know today. in this review, we provide an overview of the state-ofthe-art knowledge regarding miocene palaeogeography, vegetation and climatic conditions for the eastern north sea basin based on data from onshore denmark. this is followed by a comparison of miocene climate conditions with those of the present day for the wider north atlantic region. https://doi.org/10.34194/geusb.v57.8365 http://www.geusbulletin.org/ kasia k. śliwińska et al: geus bulletin 57. 8365. https://doi.org/10.34194/geusb.v57.8365 3 of 18 geusbulletin.org 2 danish miocene stratigraphy and the sønder (sdr.) vium core the total composite thickness of the miocene deposits onshore denmark is c. 250 m, reaching up to 400 m in local depressions. the deposits consist mainly of clay, silt and sand. the succession is subdivided into the brejning fm (partially upper oligocene), vejle fjord fm, klintinghoved fm, bastrup fm, arnum fm, odderup fm, hodde fm, ørnhøj fm, gram fm and marbæk fm (fig. 1). the succession is predominantly marine (outer to inner neritic), interfingered by fluviodeltaic deposits. each of the fluviodeltaic sequences terminates with a brown coal layer (see section 3.3 for a description). the danish miocene succession is unconformably overlain by holocene deposits. in present-day onshore denmark, the youngest miocene strata are of tortonian age (dybkjær & piasecki 2010; rasmussen et  al. 2010). the succession is well dated using dinoflagellate cysts (i.e. dinocysts; piasecki 1980; dybkjær 2004; dybkjær & piasecki 2010; rasmussen et al. 2010). a few studies have provided dinocyst stratigraphy integrated with other microfossil groups, such as foraminifera (miller et al. 1991; dybkjær & piasecki 2010; anthonissen 2012; king et al. 2016), but a complete, integrated stratigraphy for the entire danish miocene succession does not yet exist. furthermore, calcareous benthic foraminifera do not occur continuously throughout the miocene succession onshore denmark, and studies are limited to assemblage analysis (laursen & kristoffersen 1999; anthonissen 2012). currently, δ18o stratigraphy on calcareous benthic foraminifera, which could improve the stratigraphy, is lacking. despite a few attempts, the magnetostratigraphy of the miocene succession has not been resolved either and is currently limited to the oligocene–miocene boundary interval (śliwińska et al. 2014). the mco coincides with a high in global sea level (miller et al. 2020; rohling et al. 2021). in denmark, the fig. 1 lithostratigraphic framework of the uppermost oligocene–miocene onshore denmark (modified from rasmussen et al. 2010). plio.: pliocene. r. mb: resen mb. h. o.: hystrichosphaeropsis obscura. a.u.: amiculosphaera umbracula. g. v.: gramocysta verricula. a. a.: achomosphaera andalousiense. u. a.: unipontidinium aquaeductum. l. t.: labyrinthodinium truncatum. c. au.: cousteaudinium aubryae. e. i.: exochosphaeridium insigne. c. c.: cordosphaeridium cantharellus. s. h.: sumatradinium hamulatum. t. p.: thalassiphora pelagica. c. am.: caligodinium amiculum. h. spp.: homotryblium spp. c. g.: chiropteridium galea. d. p.: deflandrea phosphoritica. plio. zanclean messinian tortonian m åd e g ro up r ib e g ro up serravallian m io ce ne n eo ge ne 10 15 20 25 chattian langhian burdigalian aquitanian o lig oc en e u pp er u pp er m id dl e lo w er 5 p al ae og en e pe rio d sw ne epochma age dinocyst zonation lithostratigraphy h. o. g. v. a. a. u. a. l. t. a. u. c. c. e. i. c. au. c. g. d. p. s. h. t. p. c. am. h. spp marbæk fm gram fm ørnhøj fm hodde fm arnum fm stauning mb odderup fm bastrup fm fasterholt mb vandel mb resen mb resen mb r. mb klintinghoved fm vejle fjord fm brejning fm brejning fm skansebakke mb kolding fjord mb øksenrade mb not included in this study billund fm hvidbjerg mb addit mb sydklint mb marine silt and clay marine sand fluvial sand and gravel hiatus brackish-water silt and clay coal b aqui100 f e tort100 sequence stratigraphic subdivision lang100 burd200 d burd100 c aqui200 søby-fasterholt flora https://doi.org/10.34194/geusb.v57.8365 http://www.geusbulletin.org/ kasia k. śliwińska et al: geus bulletin 57. 8365. https://doi.org/10.34194/geusb.v57.8365 4 of 18 geusbulletin.org flooding at the base of the arnum fm is considered to be related to the onset of the mco. however, the precise onset of the mco in denmark is still uncertain. to date, the most complete miocene succession in denmark is penetrated by the sønder (sdr.) vium borehole (dgu 102.948; 55°9′04.02″ n, 8°24′46.52″e; fig. 2). the core is located in the western part of the jylland penisula, denmark, in the eastern north sea basin (fig. 3). the cored miocene succession in the sdr. vium core consists of shallow and offshore marine to deltaic deposits spanning c. 20.4–8 ma (late aquitanian – latest tortonian). the core has been studied for many years for sedimentology (rasmussen et  al. 2010), dinocyst stratigraphy (dybkjær & piasecki 2010), foraminifera stratigraphy (anthonissen 2012), pollen and spores palaeoecology, palaeotemperature and precipitation (larsson et al. 2011) and alkenone-derived sst (herbert et al. 2020). due to its exceptionally long stratigraphic range, the core has formed the backbone of the miocene stratigraphic and palaeoenvironmental studies carried out in denmark. the sdr. vium core is the only such core in the northern midand high latitudes to penetrate a near continuous miocene succession. it encompasses in ascending order: the klintinghoved fm, bastrup fm, arnum fm, odderup fm, hodde fm, ørnhøj fm and gram fm. the sequence stratigraphic framework for the core is based on the work of rasmussen (1996) and rasmussen (1997). the sequence stratigraphic framework was defined for onshore denmark and was later transferred to the danish and norwegian offshore sections (dybkjær et al. 2021). recent work has shown that the ‘d’ sequence (fig. 1) should be infact subdivided into two, referred to here as d1 (burd100) and d2 (burd200) sequences. including this subdivision allows us to correlate the sdr. vium core with the sequence stratigraphic framework for the eastern north sea basin (fig. 4), as follows: 1. sequence boundary (sb) burd100 (d1 sequence), depth 182 m based on gamma log data 2. maximum regressive surface (mrs), depth 132 m based on gamma log 3. maximum flooding surface (mfs) burd150 (d1 sequence), depth 131.75 m based on the presence of a 1 mm thick glauconite layer 4. sb burd200 (d2 sequence), depth 85 m based on gamma log data 5. mfs burd250 (d2 sequence), depth 76.5 m based on gamma log data 6. sb lang100 (e sequence), depth 51.9 m based on gamma log 7. mfs lang150 (e sequence), depth 45 m based on gamma log data fig. 2 stratigraphic framework of the sdr. vium borehole (115–24 m depth). strontium isotope ages from eidvin et al. (2014) are shown in the left column. the ages for the two uppermost samples (51.5–51.8 m and 71.15 m) show ages older than expected from the dinocyst stratigraphic framework, see discussion in eidvin et al. (2014). the neogene dinocyst zonation by dybkjær & piasecki (2010). gts 2012 from hilgen et al. (2012). first and last occurrences of age-diagnostic dinocysts are shown to the right. lithostrat.: lithostratigraphy. serravall.: serravallian. burdigal.: burdigalian. a. andalousiensis: achomosphaera andalousiensis. c.  poulsenii: cerebrocysta poulsenii. c. placacanthum: cleistosphaeridium placacanthum. c.  cantharellus: cordosphaeridium cantharellus. c. aubryae: cousteaudinium aubryae. e. insigne: exochosphaeridium insigne. g. verricula: gramocysta verricula. h. obscura: hystrichosphaeropsis obscura. p. miocaenicum: palaeocystodinium miocaenicum. u. aqueductum: unipontidinium aquaeductus. l. truncatum: labyrinthodinium truncatum. scale: 1:150 22 24 26 28 30 32 34 36 38 40 42 44 46 48 50 52 54 56 58 60 62 64 66 68 70 72 74 76 78 80 82 84 86 88 90 92 94 96 98 100 102 104 106 108 110 112 114 m ea su re d de pt h (m ) miocene p er io d/ e po ch serravall. 40.0 langhian 47.52 burdigal. 59.52 tortonian 24.0 a ge (g ts 2 01 2) ørnhøj 40.0 hodde 44.9 arnum 51.0 odderup 98.0 arnum 111.0 gram 24.0 fo rm at io n g.verriculaa.umbracula 29.14 44.12 labyrinthodinium truncatum 49.12 cousteaudinium aubryae 59.52 c. cantharellus/ e.insigne 110.77 h. obscura 24.0 zo ne 25.62 h. obscura last occurence first occurence l. truncatum 29.14 c. placacanthum 40.12 c. poulsenii 44.12g. verricula 47.52 p. miocaenicum u. aquaeductus 49.12u. aquaeductum 50.54p. miocaenicum 54.27l. truncatum common 59.52l. truncatum 59.52 c. aubryae 76.33c. poulsenii 110.77c. aubryae 110.77c. cantharellus 114.92 e. insigne chronostratigraphy lithostrat. north sea dinocyst zonation (dybkjær & piasecki 2010) e ve nt s unipontodinium aquaeductum achomosphaera andalousiense 46.62 46.62a. andalousiensis 18.44–18.38 ma 18.59–18.12 ma 18.23–17.74 ma 18.77– 17.74 ma 18.56– 17.6 ma 16.43– 16.09 ma 15.97 ma 13.82 ma 11.63 ma https://doi.org/10.34194/geusb.v57.8365 http://www.geusbulletin.org/ kasia k. śliwińska et al: geus bulletin 57. 8365. https://doi.org/10.34194/geusb.v57.8365 5 of 18 geusbulletin.org 8. sb tort100 (f sequence) at depth 40 m based on gamma log data 9. mfs tort150 (f sequence), depth c. 35 m based on clay mineralogy in the gram borehole. clay mineralogic data from sdr. vium are missing, and therefore, the position of this mfs is placed approximately. the biostratigraphic framework for the core was published by dybkjær & piasecki (2010; fig. 2), but the full set of raw data for the core were not published. in this study, we make these data available and have increased the sample resolution for the dinoflagellate dinocyst stratigraphy compared to that originally reported by dybkjær & piasecki (2010). in doing so, all slides were processed following the methods described in dybkjær (2004). the data set contains raw palynomorph counts and is provided in supplementary file s1. it includes: • 42 samples analysed in 2009 by stefan piasecki (‘sp’ in the data set) and included in dybkjær and piasecki (2010) • 10 additional samples analysed in 2022 by julie fredsborg (‘jmf’ in the data set). first and last occurrences of age-diagnostic dinocysts in the interval from 115 to 24 m are shown in fig. 2. in fig. 3 palaeogeography of the study area during the late aquitanian (a), early (b), mid (c) and late burdigalian (d), early langhian (e) and serravallian (f). white shading indicates areas where brown coal deposits were reported (after grambo-rasmussen (1984) and rasmussen et al. (2010)). purple dot: sdr. vium core. location of place names mentioned in the text is shown as black dots (cities) and white dots (localities). yellow lines show the coastline today. 100 km 100 km ed 100 km a addit brown coal 100 km100 km b c resen brown coal fasterholt brown coal silkeborg moselund addit voervadsbro dykær silkeborg sønder resen brande søby-fasterholt moselund dykær silkeborgørnhøj ørnhøj sønder resen brande dykær silkeborg sønder resen brande søby-fasterholt moselund dykær 100 km f late aquitanian (c. 21.5 ma) early burdigalian (c. 20 ma) burdigalian (c. 18.5 ma) late burdigalian (c. 16.5 ma) early langhian (c. 15.5 ma) serravallian (c. 13 ma) https://doi.org/10.34194/geusb.v57.8365 http://www.geusbulletin.org/ kasia k. śliwińska et al: geus bulletin 57. 8365. https://doi.org/10.34194/geusb.v57.8365 6 of 18 geusbulletin.org the interval from 288 to 115 m, we follow the dinocyst events as described in dybkjær & piasecki (2010). 3 the geological evolution and palaeogeography of the eastern north sea basin for more than 100 myr, from the lower cretaceous to the oligocene, the eastern north sea basin was covered by an epicontinental sea. during the oligocene, the north sea became semi-enclosed, and siliciclastics originating from present-day scandinavia were deposited. several of the cooling episodes during the oligocene are linked to falls in glacioeustatic sea level (śliwińska et al. 2010; clausen et al. 2012). at least one major fluviodeltaic system developed in the northern north sea basin during the middle oligocene (jarsve et al. 2015; śliwińska 2019). the global cooling at the oligocene–miocene boundary related to a fall in glacioeustatic sea level caused a retraction of the coastline in the area that constitutes present-day denmark. this resulted in either a hiatus or a condensed section across the boundary interval (e.g. rasmussen et al. 2010; śliwińska et al. 2014). during the early miocene, a major fluviodeltaic system developed in the eastern north sea basin (fig. 3; ziegler 1990; rasmussen et al. 2010). this change in depositional setting was caused by inversion tectonics (ziegler 1990; rasmussen et al. 2008; rasmussen 2009b). former basins within the area were uplifted, for example, the central graben and the norwegian–danish basin. the norwegian– danish basin became shallower and consequently had a physiography perfectly suited to progradation of delta systems from scandinavia and into the eastern north sea basin (olivarius et al. 2015). deposits of the fluviodeltaic systems consisted of very mature sediments, pure quartz, quartzites, kaolinite and goethite as well as flint. these types of sediments were formed due to exposure and fig. 4 temperature evolution recorded in the sdr. vium core. cmt: coldest month temperature. mat: mean annual temperature. wmt: warmest month temperature. sst: sea-surface temperature. the temperature ranges are based on the coexistence approach (ca) performed on the data set generated by larsson et al. (2010, 2011). sst based on alkenones and plotted after herbert et al. (2020). dinocyst relative abundance: relative distribution of warmand cold-favouring dinocyst taxa. dinocyst taxa assigned as warmand cold-water indicators are shown in supplementary file s2. mfs: maximum flooding surface. sb: sequence boundary. mrs: maximum regressive surface. ?mfs: possible mfs. gram tortonian (pars) stageage (ma) serravallian 11.63 13.82 15.97 20.44 ørnhøj hodde arnum arnum odderup bastrup klintinghoved langhian burdigalian water depth (m) 500 100 150 mfs lang150 mfs burd250 mfs tort150 mfs burd150 mrs sb burd100 sb burd200 sb lang100 sb tort100 ?mfs c or e de pt h (m ) 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 210 220 230 240 250 260 270 280 290 0 2 4 6 8 10 12 14 16 18 20 22 24 26 28 temperature (°c) cmt mat wmt sst dinocyst relative abundance (%) cold warm 200 40 https://doi.org/10.34194/geusb.v57.8365 http://www.geusbulletin.org/ kasia k. śliwińska et al: geus bulletin 57. 8365. https://doi.org/10.34194/geusb.v57.8365 7 of 18 geusbulletin.org erosion of old mesozoic etched basements and chalk in the hinterland (lidmar-bergström et  al. 2000). based on sediment and mineral provenance studies, the source area is known to have covered the sw part of present-day norway and central sweden (olivarius et al. 2015). 3.1 miocene shoreline the miocene landscape of the eastern north sea was characterised by a nw–se-trending shoreline (fig. 3). as a result of global sea-level change and the uplift of scandinavia (e.g. rasmussen et al. 2008), the coastline migrated ne during highstands and sw during lowstands. several times, the shoreline moved so far sw that nearly the entire area of present-day denmark could be classified as emerged lowland. in the earliest part of the miocene, the fluvial system was characterised by braided rivers but changed to meandering rivers in the later part of the early miocene (rasmussen 2009a; rasmussen et al. 2010). the eastern north sea basin was located in the northern westerly wind belt. consequently, alongshore currents transported sediments dispersed from rivers towards the se, and spit systems and barrier islands were formed se of the main delta complexes (rasmussen & bruun-petersen 2010). muddy sediments were transported nw by deeper marine bottom currents and plastered along the slope (hübscher et al. 2016). progradation and flooding of these delta systems took place three times during the early to early middle miocene. vast forested mires developed on these deltas and coastal plains, where abundant plant material accumulated, forming peats and peaty soils during each phase of delta-formation. the mires and the accumulated peat deposits were drowned and buried at the termination of each delta system, allowing the peats to be preserved as the miocene brown coal deposits we know today (figs 1, 5). the progradation and flooding of the delta systems were mainly controlled by eustatic sea-level changes (figs 1, 3; rasmussen 2004; miller et al. 2005; john et al. 2011). during the early part of the middle miocene, a major flood took place in the eastern north sea basin. the deltaand coastal plain-dominated environment of the early miocene was succeeded by fully marine depositional conditions, such that water depth in the marginal areas of the eastern north sea basin reached more than 100 m during the middle part of the middle miocene (rasmussen et al. 2010). this flooding of the delta complexes occurred despite a fall in eustatic sea level in the order of 70 m (john et al. 2011) and is interpreted to be a consequence of accelerated subsidence of the north sea basin during the middle miocene (koch 1989; rasmussen 2004). the marine conditions continued until the latest late miocene. during this period, mud deposition dominated (nielsen et  al. 2015), however, with increasing intercalation of laminated silt and fine-grained sand (rasmussen & larsen 1989). during the late miocene, the shoreline migrated across present-day denmark, and fluviodeltaic deposition was re-established (rasmussen et  al. 2005, 2010). by the end of the miocene, delta progradation commenced in the central part of the north sea basin (møller et al. 2009). 3.2 miocene water depth water depth can be reconstructed based on the fossil assemblages or established by the mapping of the clinoform break point observed in seismic profiles (rasmussen et al. 2010). the first studies providing estimates of water depth were based on molluscs (sorgenfrei 1958; rasmussen 1966, 1968). the most recent studies are based on foraminifera assemblages (laursen & kristoffersen 1999) and seismic interpretation (rasmussen et al. 2010). overall, water depths during the miocene were similar to modern-day water depths in the eastern north sea: predominantly shallow (0–50 m water depth), occasionally up to 150 m deep. the water depth estimations for the sdr. vium core are shown in fig. 4. these are based on regional correlation panels showing the architecture of the miocene succession, sequence boundaries interpreted to represent subaerial exposure (sbs burd200 and lang100), sedimentary facies shift indicating storm wave base (mfs burd250), seascape morphology from correlation panels and water depth estimation from biofacies studies (foraminifera). 3.3 miocene brown coals – a terrestrial climate archive brown coals (lignite and fossilised peat) consist of partially decayed plant material deposited in stable, wet, shallow areas, with no disturbances by rivers or oceans. areas where precursor peats formed remained saturated with water, covering dead plant material and protecting the peat from degradation by exposure to atmospheric oxygen. brown coal consists of the least altered organic material compared to other sediment types. consequently, brown coals often yield exceptionally well-preserved plant fragments, cuticles, leaves and fruits, as well as pollen and spores (terrestrial palynomorphs). in some cases, entire tree trunks or large treelog fragments have been found. fossilised peats encapsulate information about local floral diversity, temperature and precipitation. furthermore, brown coal beds provide exceptional sedimentary archives considering that they record long (thousands of years), stable depositional conditions. thus, they provide unique insights into both large-scale and local-scale changes in depositional conditions, vegetation and climate. nearly, all existing brown coal deposits suffer from a poorly constrained stratigraphic framework since https://doi.org/10.34194/geusb.v57.8365 http://www.geusbulletin.org/ kasia k. śliwińska et al: geus bulletin 57. 8365. https://doi.org/10.34194/geusb.v57.8365 8 of 18 geusbulletin.org terrestrial palynomorphs are much less accurate for dating sediments, compared with marine microfossils, and brown coal seams are primarily interbedded with sand deposits, which are also very difficult to date precisely. brown coal deposits of miocene age are known from several locations worldwide, including australia (holdgate & clarke 2000), vietnam (petersen et  al. 2022), greece and turkey (kalkreuth et al. 1991; kvaček et  al. 2002; güner et  al. 2017; bouchal et  al. 2018; oskay et  al. 2019; denk et  al. 2022) and india (singh et al. 1992). there are also well-described brown coal deposits in central and northern europe, for example, the lower rhineland in germany (utescher et al. 2000, 2002, 2021), poland (widera 2016; dumont et al. 2020; worobiec et al. 2021) and iceland (grímsson et al. 2007; denk et al. 2011). the widespread peat formation in the miocene might have been associated with high precipitation rates. for decades, brown coals were studied not only because they provide a detailed, high resolution, local terrestrial climatic record, but mainly in the context of their quality as an energy source. therefore, the best studied brown coal deposits are in countries with the largest lignite reserves and exploration potential. however, most lignite mines in europe are now closed or nearly closed as lignite is phased out to reduce global carbon emissions. brown coals were studied extensively in denmark during three exploration campaigns from the 1920s until the 1980s (koch & christensen 1979; gramborasmussen 1984). there are three brown coal-bearing layers of miocene age recognised in denmark. each of the units was deposited as a final stage of the major delta build-up episode (rasmussen et al. 2010). extensive work carried out on the miocene deposits in jylland to map aquifers means that the brown coals of denmark are relatively well dated, as follows: 1. brown coal at the top of the billund fm (the addit mb, late? middle aquitanian – equivalent to the caligodinium amiculum dinoflagellate cyst zone; rasmussen et  al. 2010, their fig. 7). the age of the zone is estimated to span an interval from ?21.6 to 21.10 ma (dybkjær & piasecki 2010). the maximum cumulative thickness of the brown coals is up to 4 m (rasmussen pers. comm) but usually 0.5 m (e.g. the voervadsbro outcrop section; rasmussen et al. 2010). petrified wood fragments were reported from the brown coal layer of this age (weibel 1996). palaeobotanical studies are limited to the lignite clays of moselund and silkeborg vesterskov (mathiesen 1965, 1970, 1975). brown coals of this age are known from the area of addit, se of silkeborg (location in fig. 3a). 2. brown coal layer at the top of the bastrup fm (within the resen mb, early burdigalian – assigned to the cordosphaeridium cantharellus zone; rasmussen et al. 2010). the stratigraphic range of the zone is from 19.0 to 18.4 ma (dybkjær & piasecki 2010). the resen mb consists of dark brown, organic-rich, silty clay with some coal layers present locally. the coaly layers are often sandwiched between sand-rich units (rasmussen et al. 2010). brown coal layers 2–3 m thick were mined close to sønder resen village between 1914 and 1970 (sdr. resen 2023). however, there are no reports of palynofloral studies of coal layers of that age. the brown coals of the resen mb are also found in southern jylland (fig. 3c). 3. brown coal at the top of the odderup fm (the fasterholt mb – assigned to the upper part of the labyrinthodinium fig. 5 the type section of the fasterholt member is the fasterholt brown coal pit, northwest of brande; this section is no longer exposed. the log is redrawn from koch (1989) and rasmussen et al. (2010). cl: clay. si: silt. f: fine sand. m: medium sand. c: coarse sand. p: pebbles. fasterholt outcrop 0 1 2 3 4 5 6 7 8 9 10 cl si f mcp sand m o dd er up f m lo w er to m id dl e m io ce ne fa st er ho lt m b sand lithology clay/silt heterolith fine medium coarse coal sedimentary structures parallel bedding planar cross-bedding biogenic structures & fossils rootlets wood after rasmussen et al. (2010) https://doi.org/10.34194/geusb.v57.8365 http://www.geusbulletin.org/ kasia k. śliwińska et al: geus bulletin 57. 8365. https://doi.org/10.34194/geusb.v57.8365 9 of 18 geusbulletin.org truncatum zone; rasmussen et al. 2010) is the thickest (fig. 5) and most spatially spread (fig. 3e). the zone is dated from 15.97 to 14.8 ma (dybkjær & piasecki 2010). the coals and coal-bearing sediments of the fasterholt mb were deposited in freshwater lakes, lagoon swamps and mires (koch 1989; rasmussen et al. 2010) and are found in several boreholes in central jylland (rasmussen et al. 2010). the cumulative thickness of the brown coals can be of up to 9 m (rasmussen pers. comm; fig. 5). the brown coal-bearing fasterholt mb is best exposed at the abildaa brown coal museum near ørnhøj. the brown coal of the fasterholt mb was extensively mined between 1914 and 1970 to supply national energy resources in denmark (e.g. gramborasmussen 1984). the flora described from this coal layer, referred to as the ‘søby-fasterholt flora’ (koch & friedrich 1970; fig. 1) or ‘damgaard flora’ (friis 1979), is so far the best studied of the miocene brown coal floras in denmark and represents vegetation during the mco (see section 4). 4 reconstructions of miocene vegetation miocene vegetation reconstructions in denmark are based on (1) palaeobotanical records comprising dispersed spores and pollen (aquitanian to tortonian, sdr. vium core; larsson et al. 2006, 2010, 2011) and (2) a rich carpological (fruits and seeds) and leaf record from the middle miocene (langhian) deposits from the søby-fasterholt flora in denmark (koch & friedrich 1970, 1971; koch et al. 1973; wagner & koch 1974; friis 1975, 1977, 1985; christensen 1976; koch 1989; denk & bouchal, 2021a, b). in this section, we outline the broad trends in vegetation development and floristic turnover during the miocene in the present-day denmark and compare these with the existing palaeobotanical records from germany and iceland (mai 1995; denk et al. 2011). while miocene floras of germany are geographically closest to denmark, the icelandic langhian to messinian macroand microfloras provide an opportunity to compare north atlantic european floras along a latitudinal gradient. during the warm humid conditions that prevailed during the early and middle miocene, coastal vegetation comprised wetlands, swamp forests and diverse, well-drained lowland forests (larsson et al. 2006, 2010, 2011). vegetation composition during this time differed considerably from the modern danish flora comprising many woody elements more typical of the present-day distribution in eastern north america and east asia (friis 1985; mai 1995; larsson et  al. 2011; denk & bouchal 2021a; supplementary file s2). the plant fossil record for the mco is derived from fruits, seeds, leaves and spores and pollen from the søby-fasterholt lignite mining area (e.g. koch & friedrich 1970; christensen 1975, 1976; friis 1985; koch 1989; denk & bouchal, 2021a, b). the record contains a high proportion of plant taxa that would be considered “exotic” today (supplementary file s3). exotic elements include genera, which are at present confined to east asia such as the conifers glyptostrobus, sciadopitys and cathaya (fig. 6a, b), the broadleaf taxa tetracentron (figs. 6c, d) and platycarya, a walnut relative. other genera are currently confined to north america, for example, the lignite forming swamp cypress taxodium and the redwood, sequoia, the woody angiosperm comptonia (fig. 6e), and wetland plants decodon (fig. 6g) and dulichium (fig. 6h). ludwigia (fig. 6j), cephalanthus (fig. 6k) and proserpinaca (fig. 6l) are also native to central and south america. today, the tree genus sequoia (redwood) is restricted to western north america. in the denmark, roots assigned to this genus were described from mco deposits (wagner & koch 1974), and leafy shoots and cones were described from premco deposits (mathiesen 1970). sequoia thrives today under a mediterranean-type climate with dry summers and moist winters. it tolerates light frost, and during the summer, it receives a lot of humidity from fog precipitation (dawson 1998). this additional supply of water during the summer suggests that the original climate niche of sequoia was fully humid temperate. in denmark, several taxa in the early and middle miocene plant assemblages represent genera that have a disjunct modern distribution in the americas and (south) east asia. among these are the conifer genus tsuga and the broadleaf taxa engelhardioideae, carya (walnut relatives), magnolia, nyssa, symplocos, liquidambar and others. as in the example of sequoia, magnolia is today found in warm temperate to tropical areas of asia and the americas but had a wide distribution in europe during the miocene. also, the wetland plant saururus (fig. 6i) belongs to this group. subtropical and tropical lineages reported from aquitanian and burdigalian pollen assemblages (e.g. the pantropical families sapotaceae and arecaceae – the palms; larsson et  al. 2006, 2011) are legacies from older periods. this is also true for several extinct taxa in the søby-fasterholt flora, for example the monocot genus cladiocarya and various extinct members of the oak family, fagaceae (daghlian 1981; friis 1985; denk & bouchal 2021a, b). a single plant genus, tetraclinis in the cupressaceae family, is today confined to sw europe and north africa. in addition to these exotic plant elements in the middle miocene of denmark, members of the modern flora of central and northern europe make up an important component (c. 1/3) of the søby-fasterholt flora fossil assemblage. these include the most characteristic tree taxa of the modern flora of denmark: pinus, fagus, alnus, https://doi.org/10.34194/geusb.v57.8365 http://www.geusbulletin.org/ kasia k. śliwińska et al: geus bulletin 57. 8365. https://doi.org/10.34194/geusb.v57.8365 10 of 18 geusbulletin.org fig. 6 typical middle miocene plant fossils from the søby-fasterholt flora lignite mining area. a–d: dispersed pollen grains. a, b: cathaya sp. (sem, lm). c, d: tetracentron sp. (sem, equatorial and polar views). e–f: leaf fossils. e: comptonia comptoniifolia (brongniart) doweld. f: acer soebyensis christensen nom. inval. g–l: dispersed fruits and seeds. g: decodon vectensis friis, seed. h: dulichium marginatum (c. reid & e.m. reid) dorofeev, fruit. i: saururus bilobatus (nikitin ex dorofeev) mai, fruitlet. j: ludwigia corneri friis, seed. k: cephalanthus pusillus friis, mericarp. l: proserpinaca brevicarpa, dorofeev endocarp. scale bars are 10 µm (a–d), 1 cm (e), 5 cm (f), 1 mm (g–i, k, l) and 0.5 mm (j). https://doi.org/10.34194/geusb.v57.8365 http://www.geusbulletin.org/ kasia k. śliwińska et al: geus bulletin 57. 8365. https://doi.org/10.34194/geusb.v57.8365 11 of 18 geusbulletin.org betula, acer and others. these taxa commonly have a disjunct modern distribution in the northern hemisphere with large distribution gaps between (eastern) north america, western eurasia and east asia. likewise, plant taxa with a cosmopolitan distribution are also common in the søby-fasterholt flora (c. 1/3 of the recorded taxa; supplementary file s4). very little is known about the pre-mco flora in denmark (e.g. larsson et al. 2011). the lignite clays of moselund and silkeborg vesterskov are associated with the progradation of the billund fm and thus are revised here to aquitanian age. leaf and wood remains as well as reproductive structures were reported from these lignites (mathiesen 1965, 1970, 1975). while the søby-fasterholt flora represents mainly riparian and wetland vegetation, the assemblage recorded from moselund and silkeborg also comprises elements from the welldrained hinterland vegetation (e.g. parrotia, daphnogene; mathiesen 1965, 1970, 1975). a marked change in the flora of denmark occurred at the onset of the late miocene (tortonian stage). investigations of dispersed spores and pollen show a sharp increase in herbaceous taxa such as the grass family (poaceae) and in ericaceae and pinus, and a decrease in engelhardioideae, extinct fagaceae (as ‘evergreen quercus’) and palms (larsson et al. 2011). the floristic composition of the miocene flora in denmark resembles, to some extent, that reported in ne germany (mai 1995). in lower and middle miocene strata, many elements are shared with the danish miocene (e.g. daphnogene, quercus sect. lobatae, mastixiaceae (syn. of nyssaceae) and palms). the palaeobotanical records from radiometrically dated strata of iceland show that middle miocene floras include exotic elements (papillate cupressaceae, cathaya, magnolia, liriodendron and others; denk et al. 2011) and lack most of the thermophilous plants present in central europe and denmark (e.g. mastixiaceae). a sharp increase in herbaceous plants is seen in the tortonian (late miocene) in denmark ( larsson et al. 2011), as well as in coeval floras of iceland, where tortonian pollen assemblages are dominated by herbaceous taxa and small-leaved ericaceae (denk et al. 2011). in contrast, mastixioid flora with lauraceae and other thermophilous elements persist into the late miocene further south in germany (mai 1995). in general, the floristic turnover from the early and middle miocene to the late miocene appears to reflect a response of the terrestrial vegetation that bordered the northern north atlantic to global cooling as indicated in the global marine isotopic records (utescher et al. 1997; westerhold et al. 2020). in contrast, the transition from late early miocene flora in denmark to mco flora, as reflected in spore and pollen assemblages (larsson et al. 2011), is not linked to marked changes in the dominant plant groups. nevertheless, there is a weak signal of warming starting in the late burdigalian and continuing into the mco. for example, sabaloid palms and mastixiaceae show increased pollen percentages during this period. the same is true for several extinct evergreen fagaceae (denk & bouchal 2021b). 5 reconstructions of air temperature and precipitation for most of the miocene, nw europe was warm and wet as suggested from records of vertebrates (mörs 2002; böhme 2003; böhme et al. 2011; utescher et al. 2011), mammals (fortelius et al. 2014), flora (mosbrugger et al. 2005; grímsson et al. 2007; bruch et al. 2011) and pollen (pound et al. 2011, 2012; pound & riding 2016; pound & mccoy 2021; mccoy et al. 2022). high levels of atmospheric moisture are also suggested by fungal assemblages (pound et al. 2022). global precipitation patterns in the miocene are characterised by large amplitudes (e.g. böhme et  al. 2011; utescher et al. 2015), and overall the mean annual precipitation (map) was 300–1000 mm higher than at present (böhme et al. 2011). in some regions, however, the climate became significantly drier. the onset of aridification in northern china is observed already in the early miocene (22 ma; e.g. guo et al. 2002). overall, aridification accelerated in the late miocene, leading to, for example, the development of the sahara desert c. 7 ma (schuster et al. 2006). aridification of the climate in the late miocene associated with the decreasing temperatures at that time was more of a global phenomenon. air temperature and precipitation from the danish miocene have been studied using the spore and pollen assemblages as well as organic molecular fossils. collectively, spore and pollen assemblages collected from two localities and one sediment core (sdr. vium) cover a time span between c. 23.5 and c. 8.5 ma, across the omt to the upper tortonian (larsson et al. 2006, 2010, 2011). notably, the data set originally generated by larsson et al. (2010, 2011) has herein been updated with the latest climate data for the nearest living relatives using the palaeoflora database and the coexistence approach (ca; utescher et al. 2014). ca analysis of the pollen record from sdr. vium indicates that the mean annual air temperature (mat) was stable during the miocene (17–18.5°c), with a mean warmest month temperature (wmt) ranging from 23 to 27.5°c, and the mean coldest month temperature (cmt) between 10 and 12.5°c (fig. 4). however, fluctuations in the miocene temperature record are inferred from expansions of mat, cmt and wmt to include colder or warmer temperatures. during the middle to late burdigalian, the ca analysis indicates that the climate included cooler temperatures, as plants requiring https://doi.org/10.34194/geusb.v57.8365 http://www.geusbulletin.org/ kasia k. śliwińska et al: geus bulletin 57. 8365. https://doi.org/10.34194/geusb.v57.8365 12 of 18 geusbulletin.org milder winter temperatures, such as mastixiaceae and sabal (palms), were absent from the pollen record. a similar pattern is observed in the serravallian and tortonian (fig. 4). during the latest burdigalian and langhian, the presence of mastixiaceae pollen, in particular, indicates slightly warmer winter temperatures (fig. 4). fluctuations in wmt correspond rather well with fluctuations in the sst (see section 6). an alternative tool for estimating mat is based on the distribution of membrane lipids derived from soil bacteria (the branched glycerol dialkyl glycerol tetraethers; brgdgt). in denmark, mat derived from brgdgt has so far been calculated only across the omt (śliwińska et al. 2014). between 24.5 and c. 22.5 ma, mat ranged from 18 to 20°c (excluding the temperature minimum of 16.5–17°c related to the cooling event at the omt). the brgdgt-derived temperature is similar (i.e. 17–20°c) to air temperature estimations based on the pollen and spores from the dykær locality, denmark (palynomorphs; larsson et al. 2010; this paper) as well as from other sites in nw europe, for example fossil leaf floras (uhl et al. 2007; uhl & herrmann 2010) and megafloras (mosbrugger et  al. 2005). however, recent progress in understanding the origin of brgdgts and improvements in calibrations for calculating palaeo-temperature call for a new analysis of the existing data set. overall, air temperature in the miocene was significantly higher than today. for comparison, mat (between 1981 and 2010) in denmark is 8.3°c (cmt = 1.0°c and occured in february; wmt = 16.6°c and occured in july; cappelen 2021). warm air temperature in central europe during the miocene is indicated by numerous published vegetation records. palynological results from germany indicate a warm temperate to subtropical climate during the miocene (zetter 1997; kolcon & sachsenhofer 1999; kovar-eder et al. 2001). palynological and palaeobotanical records from mco brown coal seams in the lower rhine basin indicate map of 1300–2000 mm and mat c. 16–20°c (utescher et al. 2021). in the lower rhine basin, wmt during the mco exceeded 25°c and sometimes reached 28°c (utescher et al. 2021). this is similar to the pollen-derived climate data from sdr. vium (fig. 4). ca analysis of the pollen record from sdr. vium indicates map c. 750–1750 mm for much of the miocene, which is much higher than the c. 749 mm observed over the past two decades in denmark. for the period 1981–2020, april is the driest month in denmark (mean monthly precipitation, mmp = 37.3 mm). the wettest month is october (mmp = 84.5 mm; cappelen 2021). overall, mmp for the driest month during the miocene ranges from c. 5 to 60 mm, while the wettest month ranges from 150 to 250 mm (fig. 7). notably, several spikes in mmp for the wettest month are observed from the late burdigalian (from c. 115 m depth in the sdr. vium core; fig. 7). both map and wettest month mmp were thus significantly higher during the mco than today, while rainfall during the driest month is within the range observed in the miocene. notably, the anomalously high rainfall of c. 150 mm reported in denmark in october 2023 reached the lower end of the reconstructed wettest month mmp for the miocene (150–250 mm). the high precipitation and warm temperatures (see also sst in section 6) in the miocene must have commonly caused storms in the eastern north sea basin. in several localities around denmark, the miocene succession yields hummocky cross-stratified strata, which suggest that intense storms were common (rasmussen et al. 2010). however, it is not possible to determine if the intensity of storms changed throughout the miocene. 6 reconstructions of sst the sst estimations from the sdr. vium core suggest temperatures of up to 28°c, 20.4–12 ma (early to middle miocene), followed by a gradual decrease to c. 24°c, 12–8 ma (late miocene; fig.  4; herbert et  al. 2020). here (fig. 8), we follow the age model (absolute ages) proposed by herbert et al. (2020). sst is derived from alkenone palaeothermometry (uk 37 index defined by prahl & wakeham 1987). overall, temperature trends for the sdr. vium core agree with sst records from six other sites in the north atlantic region (fig. 8). however, on closer inspection, pre-mco temperatures were already high (25–28°c), with no apparent trend in time. notably, just before the onset of the mco, a brief interval of lower temperatures (i.e. the sst minimum; 24–25.5°c) marks the most pronounced feature of the sst trend for denmark. this sst minimum is observed at sites 608, 982 and the sdr. vium core – the only sites covering the mco onset. no major increase in sst related to the onset of the mco was observed at any of these three sites (fig. 8). discrepancies in the timing of the sst minimum may be related either to uncertainties in the age models (especially for the sdr. vium core; see herbert et al. 2020 for details) or because these sites are in a region of dynamic oceanic circulation intensity and at the northern extent of the subtropical gyre. overall, an improved high resolution stratigraphic framework (magnetoor isotope stratigraphy) would significantly aid in explaining these trends. we could speculate that the sst minimum is related to a colder interval just prior to the mco, as observed in the sedimentary records from offshore antarctica (levy et  al. 2016). however, we observe that the sst minimum is observed only in the sites bathed by the modern-day subtropical gyre, while site u1406, which is under a subpolar gyre regime, shows a gradual temperature increase in the time-equivalent interval. we https://doi.org/10.34194/geusb.v57.8365 http://www.geusbulletin.org/ kasia k. śliwińska et al: geus bulletin 57. 8365. https://doi.org/10.34194/geusb.v57.8365 13 of 18 geusbulletin.org could, therefore, argue that the sst minimum observed in the eastern north atlantic associated with warming in the western north atlantic (reorganisation of the subtropical gyre) was an effect of weakening of the oceanic circulation in the north atlantic rather than an overall cooling. however, this would need to be supported by, for example, modelling. 7 the onset of mco and the middle to late miocene transition following the age model applied by herbert et al. (2020), the onset of the mco is placed at the onset of warmer sst (approximately at depth 90 m in the sdr. vium core). however, both dinocyst assemblages (showing a flux of warm water taxa) and pollen and spore assemblages suggest that warming occurred earlier and stepwise. the relative abundance of warm water dinocysts shows an initial increase at c. 130 m depth and another increase at 110 m (fig. 3). the first flux of warm water dinocysts at 130 m corresponds with the mfs at the base of the arnum fm, which is considered to be synchronous with the global sea-level rise at the onset of mco. however, sst derived from alkenones record a sharp decrease at that depth, which contradicts the dinocyst records. our climate inferences based on the palynological record suggest at least three warming steps: (1) at depth 110 m, associated with a slight increase in the relative abundance of palm pollen, (2) at 90 m, where the abundance of warmth-loving deciduous fagaceae (‘castanea’ in larsson et  al. 2011) increased, and (3) at 75 m, where sabaloid palms, engelhardioideae and the evergreen fagaceae showed a marked increase. the transition from the middle to the late miocene is well expressed in our inferred palaeoclimate data fig. 7 precipitation trends derived from coexistence approach (ca) on the updated pollen and spore record from the sdr. vium core, generated by larsson et al. (2010, 2011). dark green triangle: mean annual precipitation (map) in denmark, 1981–2020 (c. 749 mm; cappelen 2021). lightest green triangle: mean monthly precipitation (mmp) for the driest month (april; 37.3 mm). light green triangle: mmp for the wettest month (october; 84.5 mm). 30 depth [m] 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 210 220 230 240 250 260 270 280 290 [mm]0 200 400 600 800 1000 1200 1400 1600 1800 gramtortonian (pars) stage formationage (ma) serravallian 11.63 13.82 15.97 20.44 ørnhøj hodde arnum odderup bastrup klintinghoved langhian burdigalian map mmp wettest month mmp driest month https://doi.org/10.34194/geusb.v57.8365 http://www.geusbulletin.org/ kasia k. śliwińska et al: geus bulletin 57. 8365. https://doi.org/10.34194/geusb.v57.8365 14 of 18 geusbulletin.org and in corresponding species turnover in the terrestrial and marine realms. in the plant fossil record, this transition is reflected in a marked decrease of taxodiaceous cupressaceae, and an increase in cold-tolerant woody taxa, such as carya, ericaceae and poaceae. likewise, the post-mco cooling trend in surface water temperature is recorded by the alkenones and the sharp increase in cold-tolerant dinocysts. 8 conclusions the mco is the most recent time in earth history where the co2 levels were as high as today. all the existing data from the danish miocene succession suggest that the area was significantly warmer and wetter than today. however, recent anomalies in rainfall (such as the massive rainfall in october 2023) reach the minimum values recorded for mmp during the wettest month in the miocene. the onset of the mco in the eastern north sea basin seems to be related to a sequence of warming steps or pulses in both the marine and terrestrial realms. some of these steps are synchronous, some are not. however, the main obstacle in the evaluation of these is the fact that the existing data sets are either of too low resolution or have different sampling depths. another improvement in resolving the miocene climate record for the eastern north sea basin could be to provide an integrated magnetostratigraphic age model for the danish miocene. acknowledgements we thank diederik liebrand for preparing fig. 8. we would like to dedicate this paper to dr linda larsson who generated the pollen and spore record during her phd project and who passed away in 2023. additional information author contributions ks: designed the research, generated original figs 2, 4, 7, and wrote the paper with contribution from kd, esr, sl and td. td: wrote the section about miocene vegetation, prepared fig. 6, and compiled the data set in s4. sl: compiled and updated the data sets from larsson et al. (2006; 2011). sp: analysed dinocysts from 42 samples from the sdr. vium core. jmf: analysed dinocysts from 10 samples from the sdr. vium core. fig. 8 compilation of existing sea-surface temperature (sst) records from the northern low to high latitudes. sites are as follows: deep sea drilling program (dsdp) site 608, ocean drilling program (odp) site 907, odp site 982, integrated ocean drilling program (iodp) site u1318, iodp site u1404, iodp site u1406 and sdr. vium core. data sources: herbert et al. (2016, 2020); liu et al. (2018); super et al. (2018, 2020); guitián et al. (2019); sangiorgi et al. (2021). also shown is the global composite oxygen isotope curve (δ18o) for benthic foraminifera (westerhold et al. 2020) across the miocene. mco: miocene climatic optimum. cha.: chattian. oli.: oligiocene. plio.: pliocene. zan.: zanclean. c3–c6: polarity chrones. uk 37: unsaturated long-chain ketones index with 37 carbon atoms (c37 alkenones). tex86: tetraether index of 86 carbon atoms. 907 608 sdr. vium 982 u1318 u1404 u1406 pli. miocene oli. zan. messinian tortonian serravallian langhian burdigalian aquitanian cha. c3ac3 .b c4 c4a c5 c5a c5b c5c c5d .5e c6 c6a c6b c6c 18 o (‰ ) 1 2 3 4 s s t (° c ) 0 10 20 30 40 age (ma) 4 6 8 10 12 14 16 18 20 22 24 907 uk 37 982 uk 37 982 tex86 608 uk 37 608 tex86 u1318 uk 37 u1318 tex86 u1406 uk 37 u1406 tex86 sønder vium uk 37 u1404 uk 37 duration of sst min. mco https://doi.org/10.34194/geusb.v57.8365 http://www.geusbulletin.org/ kasia k. śliwińska et al: geus bulletin 57. 8365. https://doi.org/10.34194/geusb.v57.8365 15 of 18 geusbulletin.org funding statement this work has been funded by research grant no. 2-2019 awarded by geocenter denmark to kks and by research grant no. 2021-05849 from the swedish research council to td. competing interests the authors declare no competing interests. additional files four supplementary files are available at https://doi.org/10.22008/ fk2/t1yjc4. supplementary file s1: raw polymorph counts. supplementary file s2: full data set of dinocyst raw counts plotted using stratabugs program. supplementary file s3: raw pollen counts. supplementary file s4: søby fasterholt and moselund silkeborg floras. supplementary s5: sdr. vium ca-derived estimations for precipitation and temperature. references anagnostou, e., john, e.h., edgar, k.m., foster, g.l., ridgwell, a., inglis, g.n., pancost, r.d., lunt, d.j. & pearson, p.n. 2016: changing atmospheric co2 concentration was the primary driver of early cenozoic climate. nature 533, 380–384. https://doi.org/10.1038/nature17423 anthonissen, e.d. 2012: a new miocene biostratigraphy for the northeastern north atlantic: an integrated foraminiferal, bolboformid, dinoflagellate and diatom zonation. newsletters on stratigraphy 45, 281–307. https://doi.org/10.1127/0078-0421/2012/0025 badgley, c., barry, j.c., morgan, m.e., nelson, s.v., behrensmeyer, a.k., cerling, t.e. & pilbeam, d. 2008: ecological changes in miocene mammalian record show impact of prolonged climatic forcing. proceedings of the national academy of sciences 105, 12145–12149. https:// doi.org/10.1073/pnas.0805592105 böhme, m. 2003: the miocene climatic optimum: evidence from ectothermic vertebrates of central europe. palaeogeography, palaeoclimatology, palaeoecology 195, 389–401. https://doi.org/10.1016/ s0031-0182(03)00367-5 böhme, m., winklhofer, m. & ilg, a. 2011: miocene precipitation in europe: temporal trends and spatial gradients. palaeogeography, palaeoclimatology, palaeoecology 304, 212–218. https://doi.org/10.1016/j.palaeo.2010.09.028 bouchal, j.m., güner, t.h. & denk, t. 2018: middle miocene climate of southwestern anatolia from multiple botanical proxies. climate of the past 14, 1427–1440. https://doi.org/10.5194/cp-14-1427-2018 bruch, a.a., utescher, t. & mosbrugger, v. 2011: precipitation patterns in the miocene of central europe and the development of continentality. palaeogeography, palaeoclimatology, palaeoecology 304, 202–211. https://doi.org/10.1016/j.palaeo.2010.10.002 cappelen, j. 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lithostratigraphic framework of the uppermost oligocene–miocene onshore denmark (modified from rasmussen et al. 2010). plio.: pliocene. r. mb: resen mb. h. o.: hystrichosphaeropsis obscura. a.u.: amiculosphaera umbracula. g. v.: gramocysta verricula. a. a.: achomosphaera andalousiense. u. a.: unipontidinium aquaeductum. l. t.: labyrinthodinium truncatum. c. au.: cousteaudinium aubryae. e. i.: exochosphaeridium insigne. c. c.: cordosphaeridium cantharellus. s. h.: sumatradinium hamulatum. t. p.: thalassiphora pelagica. c. am.: caligodinium amiculum. h. spp.: homotryblium spp. c. g.: chiropteridium galea. d. p.: deflandrea phosphoritica. fig. 2 stratigraphic framework of the sdr. vium borehole (115–24 m depth). strontium isotope ages from eidvin et al. (2014) are shown in the left col-umn. the ages for the two uppermost samples (51.5–51.8 m and 71.15 m) show ages older than expected from the dinocyst stratigraphic frame¬work, see discussion in eidvin et al. (2014). the neogene dinocyst zona¬tion by dybkjær & piasecki (2010). gts 2012 from hilgen et al. (2012). first and last occurrences of age-diagnostic dinocysts are shown to the right. lithostrat.: lithostratigraphy. serravall.: serravallian. burdigal.: bur¬digalian. a. andalousiensis: achomosphaera andalousiensis. c. poulsenii: cerebrocysta poulsenii. c. placacanthum: cleistosphaeridium placacanthum. c. cantharellus: cordosphaeridium cantharellus. c. aubryae: cousteaudin¬ium aubryae. e. insigne: exochosphaeridium insigne. g. verricula: gramo¬cysta verricula. h. obscura: hystrichosphaeropsis obscura. p. miocaenicum: palaeocystodinium miocaenicum. u. aqueductum: unipontidinium aquae¬ductus. l. truncatum: labyrinthodinium truncatum. fig. 3 palaeogeography of the study area during the late aquitanian (a), early (b), mid (c) and late burdigalian (d), early langhian (e) and serravallian (f). white shading indicates areas where brown coal deposits were reported (after grambo-rasmussen (1984) and rasmussen et al. (2010)). purple dot: sdr. vium core. location of place names mentioned in the text is shown as black dots (cities) and white dots (localities). yellow lines show the coastline today. fig. 4 temperature evolution recorded in the sdr. vium core. cmt: coldest month temperature. mat: mean annual temperature. wmt: warmest month temperature. sst: sea-surface temperature. the temperature ranges are based on the coexistence approach (ca) performed on the data set generated by larsson et al. (2010, 2011). sst based on alkenones and plotted after herbert et al. (2020). dinocyst relative abundance: relative distribution fig. 5 the type section of the fasterholt member is the fasterholt brown coal pit, northwest of brande; this section is no longer exposed. the log is redrawn from koch (1989) and rasmussen et al. (2010). cl: clay. si: silt. f: fine sand. m: medium sand. c: coarse sand. p: pebbles. fig. 6 typical middle miocene plant fossils from the søby-fasterholt flora lignite mining area. a–d: dispersed pollen grains. a, b: cathaya sp. (sem, lm). c, d: tetracentron sp. (sem, equatorial and polar views). e–f: leaf fossils. e: comptonia comptoniifolia (brongniart) doweld. f: acer soebyensis christensen fig. 7 precipitation trends derived from coexistence approach (ca) on the updated pollen and spore record from the sdr. vium core, generated by larsson et al. (2010, 2011). dark green triangle: mean annual precipitation (map) in denmark, 1981–2020 (c. 749 mm; cappelen 2021). lightest green triangle: mean monthly precipitation (mmp) for the driest month (april; 37.3 mm). light green triangle: mmp for the wettest month (october; 84.5 mm). fig. 8 compilation of existing sea-surface temperature (sst) records from the northern low to high latitudes. sites are as follows: deep sea drilling program (dsdp) site 608, ocean drilling program (odp) site 907, odp site 982, integrated ocean drilling program (iodp) site u1318, iodp site u1404, iodp site u1406 and sdr. vium core. data sources: herbert et al. (2016, 2020); liu et al. (2018); super et al. (2018, 2020); guitián et al. (2019); sangiorgi et al. (2021). also shown is the global composite oxygen isotope curve (δ18o) for benthic foraminifera (westerhold et al. 2020) across the miocene. mco: miocene climatic optimum. cha.: chattian. oli.: oligiocene. plio.: pliocene. zan.: zanclean. c3–c6: polarity chrones. uk37: unsaturated long-chain ketones index with 37 carbon atoms (c37 alkenones). tex86: tetraether index of 86 carbon atoms. schovsbo et al. 2025: geus bulletin 60. 8383. https://doi.org/10.34194/62417j08 1 of 17 research article assessment of formation brine salinity, pressure and temperature in selected structures in eastern denmark and implications for co2 storage niels h. schovsbo*1 , hanne d. holmslykke2 , anders mathiesen3 , carsten m. nielsen1 1department of geo-energy and storage, geological survey of denmark and greenland (geus), copenhagen, denmark; 2department of geochemistry, geological survey of denmark and greenland (geus), copenhagen, denmark; 3department of geophysics and sedimentary basins, geological survey of denmark and greenland (geus), copenhagen, denmark abstract co2 storage presents new risks and challenges, where the properties of formation water play an important role. these challenges include reduced injectivity and storage capacity due to salt precipitation, viscous fingering caused by viscosity contrasts between co2 and brine and diminished co2 solubility in formation waters. understanding these factors and developing predictive models for pressure distribution are essential for successful co2 storage projects. this study presents salinity (cl and total dissolved solids), density, temperature, pressure, halite (nacl) saturation, co2 solubility and viscosity of formation waters across five co2 storage sites in denmark (stenlille, gassum, rødby, lisa and inez), covering eight reservoirs (one in the frederikshavn formation, four in the gassum formation and three in the bunter sandstone and skagerrak formations). salinity assessments are based on existing brine data or, where unavailable, a reference salinity model developed from a water chemistry database with 77 analyses from 28 wells in the danish basin and adjacent regions. the model was created using partial least squares regression, accounting for local geological developments and subsurface salts. we report high chloride levels (182 000–202 000 mg/l) and densities (1.21–1.23 kg/l) in the bunter sandstone and skagerrak formations, while the gassum and frederikshavn formations are undersaturated with halite, exhibiting lower chloride levels (99 000–148 000 mg/l) and densities (1.11–1.17 kg/l). these differences suggest a higher risk of mineral precipitation due to brine evaporation in dry co2, and a higher risk of density override due to significant density contrast, which will hamper filling efficiency in older reservoirs. modelling shows that co2 solubility reaches 33.9 g co2/l, with a 37% reduction due to chemical and pressure– temperature variations. conceptual fluid flow modelling is recommended to further assess brine– rock–co2 interactions. the salinity model has implications for geothermal reservoir assessment and can be applied regionally. *correspondence: nsc@geus.dk received: 06 oct 2024 revised: 17 dec 2024 accepted: 27 jan 2025 published: 21 may 2025 keywords: brine–co2 dynamics, co2 solubility, fluid–rock interaction, geochemical modelling, halite saturation, viscous fingering abbreviations: geus: geological survey of denmark and greenland pls: partial least squares pw: hydrostatic pressure scco2: supercritical co2 si: saturation index srm: spatial relationship matrix t: temperature tds: total dissolved solids δw: brine density μ: viscosity geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: jon ineson (geus, denmark) reviewed by: philip ringrose (ntnu, norway), kim h. esbensen (khe consulting, denmark) funding: see page 15 competing interests: see page 15 additional files: see page 15 1. introduction deep underground storage of co2 in suitable saline aquifers or depleted oil and gas reservoirs represents a significant strategy to mitigate global warming caused by high levels of anthropogenic co2 emissions (ipcc 2023). in denmark, there are many porous rock formations, which offer a high technical co2 storage potential, with multiple large structures suitable for storage of dense, supercritical (sc)co2 (hjelm et  al. 2022; gregersen et al. 2025, this volume). over recent years, characterising these sites to advance their development into co2 storage facilities has been a major focus for the geological survey of denmark and greenland (geus; hjelm et al. 2022; gregersen et al. 2023, 2025; abramowitz et al. 2024; bjerager et al. 2024; fyhn et al. 2024; keiding et al. 2024). the primary reservoirs under investigation onshore denmark and in the eastern part of the offshore sector (fig. 1) are triassic to jurassic sandstones (bunter sandstone, skagerrak and gassum formations, fig. 2) that form https://doi.org/10.34194/62417j08 https://orcid.org/0000-0003-4723-0586 https://orcid.org/0000-0002-5139-4033 https://orcid.org/0000-0003-4345-7513 https://orcid.org/0000-0002-1525-1385 mailto:nsc@geus.dk https://creativecommons.org/licenses/by/4.0/deed.ast schovsbo et al. 2025: geus bulletin 60. 8383. https://doi.org/10.34194/62417j08 2 of 17 geusbulletin.org regional aquifers. the presence of high salinity formation water in these aquifers has been well-documented since the 1940s, marking the beginning of hydrocarbon exploration activities. an initial survey of formation water chemistry was conducted by dinesen (1961), which was further elaborated by laier (1989a, 1989b, 2002, 2008). traditionally, the discussion around formation water chemistry has been focused on implications for geothermal energy production, particularly concerning the risks associated with steel corrosion in highly saline environments and the potential for scale and mineral precipitation, which can severely impact operational efficiency and safety (laier 2002; holmslykke et al. 2019, 2023; kazmierczak et al. 2022). with the resurgence of interest in the triassic and jurassic reservoirs for co2 storage applications, the need for pre-drilling assessments of formation water characteristics for specific structures, along with regional mapping of water properties such as density, has gained prominence. moreover, co2 storage introduces novel risks and challenges where formation water plays a key role. these include the risk of injection and storage capacity reduction due to salt precipitation as a result of reservoir dry-out (pruess 2009; ringrose 2020; edem et  al. 2022; cui et  al. 2023; worden 2024) and complications from viscous fingering arising from significant density and viscosity contrasts between brine and scco2, which may impede the effective distribution of co2 throughout the reservoir (kumar et al. 2020; ringrose 2020). additionally, the solubility of co2 in formation water decreases as salinity increases, which can significantly diminish co2 dissolution and, consequently, the overall efficiency of co2 sequestration (holloway 2005; deng et  al. 2018). in addition, maintaining pressure control in the reservoir by discharging formation water may be a key component for maintaining caprock and reservoir integrity (dewar et al. 2022). however, the environmental impact arising from such discharge must be assessed, including the possible mixing ratio with seawater, as even a slight increase in salinity compared to ambient levels is regarded as harmful to the aquatic environment. understanding these interactions and developing accurate predictive models are essential for the design and successful implementation of co2 storage projects. in this study, we use existing databases to conduct assessments of various mesozoic structures that formed part of the geus-led carbon capture and storage study, ccs2022–2024, focused on maturing potential co2 storage sites (gregersen et al. 2023, 2025; fig. 1). the objective is to assess the key physical properties of the formation water including its temperature, pressure, salinity (cl and total dissolved solids (tds)) and density, to discuss the effects hereof on the co2 storage efficiency and hence to highlight injection risks. in order to evaluate salinities in structures and reservoir levels where no water analyses fig. 1 location of wells and potential co2 storage sites (structures) mentioned in the text. well name abbreviations: er-4s: erslev-4s. fa-1: farsø-1. ga-1: gassum-1. ha-1: haldager-1. hö-2: höllviken-2. hö-1: höllviksnäs-1. in-1: inez-1. fcc-1: malmö/fcc-1. ma-1: margretheholm-1. ma-2: margretheholm-2. rø-1: rødby-1. st-1–6: stenlille-1, stenlille-2, stenlille-3 stenlille-4 stenlille-5 and stenlille-6. st-19: stenlille-19. stv-1: stevns-1. sø-1a: sønderborg-1a. sø-2: sønderborg-2. th-2: thisted-2. th-3: thisted-3. tø-1: tønder-1. tø-4: tønder-4. tø-5: tønder-5. ør-1: ørslev-1. aa-1a: aars-1a. aar-02: aarhus-02. 14°e13°e12°e11°e10°e9°e8°e7°e6°e 57°n 56°n 55°n legend well with formation water chemistry danish exploration well assessed salt structures major faults structural highs salt structures er-4s fa-1 aar-02 eg(s)-1 ga-1 ha-1 hö-2 hö-1 fcc-1 dge-1ma-1 ma-2 st-1–6 st-19 stv-1 sø-1a sø-2 th-2 th-3 tø-1 tø-4tø-5 ør-1 aa-1a j-1x in-1 inez lisa rødby rø-1 gassum stenlille danish basin german basin ringkøbing fyn high sjællandjylland 100 km https://doi.org/10.34194/62417j08 http://www.geusbulletin.org/ schovsbo et al. 2025: geus bulletin 60. 8383. https://doi.org/10.34194/62417j08 3 of 17 geusbulletin.org fig. 2 mesozoic stratigraphy of the studied area. reproduced from gregersen et al. (2025, this volume). röt fm bunter sandstone fm bunter shale fm zechstein kupferschiefer muschelkalk fm keuper fm 145 150 155 160 165 170 175 180 185 190 195 200 140 135 80 85 90 95 100 105 110 115 120 125 130 75 70 260 265 240 245 250 255 235 230 220 225 210 215 205 u pp er ju ra ss ic m id dl e ju ra ss ic lo w er ju ra ss ic tr ia ss ic pe rm ia n lo w er c re ta ce ou s u pp er c re ta ce ou s oxfordian kimmeridgian volgian callovian bathonian bajocian aalenian toarcian pliensbachian sinemurian hettangian rhaetian norian carnian ladinian anisian olenekian induan changhsingian capitanan wordian wuchiapingian maastrichtian campanian santonian coniacian turonian cenomanian albian aptian barremian hauterivian valanginian ryazanian skagerrak fm bunter sandstone fm bunter shale fm chalk gp rødby fm vedsted fm børglum fm flyvbjerg fm fjerritslev fm mcu ecu gassum fm vinding fm oddesund fm falster fm ørslev fm zechstein gp rotliegend gp haldager sand fm frederikshavn fm tønder fm hiatus arid lacustrine, restricted marine mudstones, evaporites and thin sandstones fluvial and aeolian sandstones and mudstones marine marlstones and mudstones marine limestones and mudstones evaporites and mudstones shallow marine and �uvial sandstones marine mudstones marine chalk 65 paleocene danian ngb rfh db stz ns rø db y2 sø lle st ed -1 ø rs le v1 sl ag el se -1 st en lil le -1 9 u lle rs le v1 t e rn e1 u. permian – triassic lithostratigraphy of northern germany mcu: mid-cimmerian unconformity ecu: early cimmerian unconformity https://doi.org/10.34194/62417j08 http://www.geusbulletin.org/ schovsbo et al. 2025: geus bulletin 60. 8383. https://doi.org/10.34194/62417j08 4 of 17 geusbulletin.org are available, we have developed a methodology based on the well-documented depth dependencies identified in previous research (laier 1989a, 2008; holmslykke et al. 2019) that aims to refine our predictions by aligning them more closely with the specifics of the local geology using the methodology presented in schovsbo et al. (2020). we use the term ‘reservoir’ here as an informal zone of a lithostratigraphical unit to describe a porous unit suitable for co2 storage. 2. geological setting gregersen et  al. (2025, this volume) have recently described the geological development including basin tectonics and history of the danish basin and north german basin and the reader is referred to this publication for a detailed account. here, we provide a more focused description highlighting the geological and deposition characteristics that are most relevant for the evaluation of formation water salinity. the structural configurations of the danish basin and the north german basin (fig. 1) were established during the early permian (michelsen et al. 2003). in the central part of the danish basin, the post mid-permian succession attains thicknesses up to 9000 m (fig. 3d). the upper permian succession (zechstein group) is characterised by extensive evaporite deposits, particularly in the major depocentres, with carbonate upon and fringing paleo highs (ziegler 1990; geluk 2000; fig. 3a). evaporite accumulation waned in the earliest triassic, when sedimentation of non-marine siliciclastics dominated the area, but resumed in the late early triassic. in the north german basin, evaporites accumulated during the extensional tectonics of the hardegsen phase, forming the röt basin in the latest early triassic (kovalevych et  al. 2002; fig.  3b). in denmark, correlative deposits are represented by the ørslev formation (fig. 2), an evaporitic claystone unit that includes rock salt (mostly halite) deposits in the deeper parts of the basin (bertelsen 1980). the northern limit of the halite facies in the ørslev formation towards the ringkøbing– fyn high is somewhat uncertain but has been re-evaluated based on well logs and cuttings during  the preparation of the geothermal webgis portal and likely extend north up to the ringkøbing fyn high (see vosgerau et al. 2016; fig. 3b). evaporite accumulation resumed in the late triassic. although the upper triassic in the danish basin and the north german basin is predominately characterised by terrestrial siliciclastics, evaporitic conditions prevailed locally in the former basin during deposition of the oddesund formation (bertelsen 1980). this is typically an evaporite-bearing claystone succession, but significant rock salt deposits have been detected in central depocentres of the danish basin (fig. 3c). additional mapping associated with the ccs2022–2024 project has shown that oddesund evaporites are not limited to these areas but may have formed in paleo lows elsewhere in the basin (fyhn et al. 2024). in the context of co2 storage potential, the key mesozoic reservoirs in the danish subsurface are the braided river and aeolian sandstones of the lower to mid triassic bunter sandstone and skagerrak formations, deposited under arid to semi-arid conditions, and the upper triassic – lower jurassic gassum and the cretaceous frederikshavn formations, deposited in paralic and shallow marine environments (fig. 2). the bunter sandstone formation is the main triassic reservoir unit in the north german basin and the danish basin, reaching thicknesses of up to 300 m (bertelsen 1980); within the danish basin, this formation grades laterally into the temporally equivalent lower levels of the skagerrak formation at the basin margin (bertelsen 1980; nielsen 2003; weibel & friis 2004; gregersen et al. 2025; keiding et al. 2024). the gassum formation represents the main reservoir sandstone unit in the area; it exhibits thicknesses ranging from 50 to 150 m across both the north german basin and the norwegian–danish basin, though it is absent over the ringkøbing–fyn high and thins into the fennoscandian border zone. marine conditions were established from the late triassic through the early jurassic, the transgressive development being recorded by the restricted marine vinding formation, the fluvial to shallow marine gassum formation and the offshore marine fjerritslev formation (fig. 2). mid-jurassic uplift led to the erosion and truncation of the fjerritslev formation. subsequent marine deposition included the accumulation of additional secondary sand-rich reservoirs such as the middle jurassic haldager sand formation and the lower cretaceous frederikshavn formation. 3. data and methods 3.1. geochemical database we use the geochemical database of formation water compositions originally compiled by laier (1989a, 2008), now augmented with new data from holmslykke et  al. (2019) covering the sønderborg, thisted and margretheholm geothermal test sites. additional contributions include data from laier (1989b) from the stenlille area, from bonnesen et al. (2009) concerning the deepest section of the stevns-1 research well (at 450 m) to limit the contributions of meteoric water in the shallower parts (bonnesen et al. 2009) and a zechstein group water sample from the ørslev-1 well (gulf denmark 1968, p.  36). data also comes from the https://doi.org/10.34194/62417j08 http://www.geusbulletin.org/ schovsbo et al. 2025: geus bulletin 60. 8383. https://doi.org/10.34194/62417j08 5 of 17 geusbulletin.org geothermal well dge-1 that tested the basement in scania (rosberg & erlström 2022) and from the geothermal well aarhus-02 that tested the gassum formation between 2293 and 2411 m (fig. 1). we excluded data from the topmost section (above 200 m) of the shallow erslev-4s well due to the high likelihood of groundwater contamination (laier 1989a; bonnesen et al. 2009). the database now encompasses 28 wells (fig. 1) and 77 brine analyses (detailed in supplementary table s1). as noted by laier (1989a, 2008), data on formation water composition are derived from two principal sources: well tests conducted during exploration or production tests conducted at geothermal plants or natural gas storage sites. the first source often presents significant uncertainties regarding contamination from the drilling process, in contrast to the latter, which offers more reliable data but with considerably more limited geographical coverage. consistent with approaches in previous data compilations, our study incorporates both types of data, including brine extracted from core samples (supplementary table s1). our analytical data span more than 75 years of laboratory and methodological development, fig. 3 maps showing extension of rock salt in (a) the zechstein, (b) the röt (ørslev formation), and (c) the oddesund formation and (d) the depth to the pre zechstein surface. a is modified after geluk (2000), b after kovalevych et al. (2002) and vosgerau et al. 2016 (webgis röt salt risk map that highlights a high-risk area for rock salt presence), c after bertelsen (1980) and d after vejbæk & britze (1994). !(!(!( !(!(!(!( !(!( !( !(!( !( !( !(!(!( !(!(!(!(!( !(!( !( !( !( d 0–1 1–2 2–3 3–4 4–5 5–6 6–7 7–8 8–9 9–10 depth (km) zechstein platform zechstein salt basina b röt salt röt salt risk (webgis) c oddesund salt 100 km https://doi.org/10.34194/62417j08 http://www.geusbulletin.org/ schovsbo et al. 2025: geus bulletin 60. 8383. https://doi.org/10.34194/62417j08 6 of 17 geusbulletin.org introducing a layer of uncertainty to our analysis. however, we have focused on the main conservative element, notably chloride (cl), which mitigates some of the potential issues related to data variability over time. the charge balance of the water analysis containing all major ions was calculated using the software programme phreeqcv3 (parkhurst & appelo 2013). the results indicate charge balance errors less than ±3% with a few samples with a charge balance error of up to –5.6% (supplementary table s1). charge balances of up to 5% are typically deemed acceptable for diluted samples (appelo & postma 2005). thus, the data on the formation water composition were presumed robust considering the high salinity of the samples. 3.2. methods 3.2.1. estimation of total dissolved solids the total dissolved ions in a sample represent the sum of all analysed ions. for samples without a full composition analysis, we use an empirical relationship between chloride and tds. the relationship is established from the mesozoic reservoirs in fig. 4 and is given by: tds (mg/l) = 1.656 · cl (mg/l) – 3528 (mg/l) (1) in eq. (1), the empirically determined coefficient 1.656 is slightly higher than that expected for a pure nacl solution (1.649), indicating that while nacl is the predominant ion pair, other ions such as ca and k are also present (see also holmslykke et al. 2019). 3.2.2. brine density, halite saturation and co2 solubility using the computer software programme phreeqcv3 (parkhurst & appelo 2013) and its pitzer database, we calculated the brine density (δw) and halite (nacl) saturation index (sinacl) at both ambient and reservoir conditions (i.e. at in situ pressure and temperature), the co2 solubility at reservoir conditions using the peng–robinson equation of state and the density of fully co2-saturated formation water (δwco2) at reservoir conditions. due to the interactions of different ions, this could only be performed for samples with a full composition analysis (table 2 and 3). for samples with measured chloride concentrations but with no full compositional analysis, we used eq. (1) to calculate tds and then ambient δw from collins (1987) using eq. (2): δw (kg/l) = 0.695 · tds (kg/l) + 1.0 (kg/l) (2) comparison between δw calculated from a full composition analysis using phreeqc with δw calculated using eq. (2) shows similar results with a deviation of only 0.01 kg/l between the two estimates. the halite saturation state of the brines is indicated by the sinacl whereby positive and negative values indicate super-saturation and undersaturation, respectively. formation water with a sinacl value within the range –0.4 ≤  sinacl ≤ +0.4 is assumed to be saturated, and thus in equilibrium with halite. the range accounts for the uncertainties associated with the difficulties of sampling brines at high temperature and pressure, the analytical uncertainty and the application of thermodynamic equilibrium constants on mineral phases in saline systems (holmslykke et al. 2019). the sinacl value could only be calculated for samples in which both na and cl were measured. 3.2.3. hydrostatic pressure and temperature the danish basin and the danish part of the german basin are considered normally pressured based on experience gained from drilling activities, and the hydrostatic pressure (pw) is calculated using geus’ own regional pressure model: pw (bar) = 0.1054 (bar/m) · [depth in m below ground level] (3) where 0.1054 bar/m is the regional pressure gradient. the gradient is often assumed to be invariant, but in practice, the gradient depends on the brine salinity, which is known to be a function of depth (laier 1989a). hence the gradient in eq. (3) should be viewed as an average of the changing brine densities in the gassum formation. validation of the pressure model with in situ fig. 4 empirical relationships between chloride and tds concentrations established for mesozoic reservoirs. based on data in laier (2008) and holmslykke et al. (2019). cl (mg/l) 1.0e+5 1.5e+5 2.0e+5 to ta l d is so lv ed s ol id s (t d s m g/ l) 1.5e+5 2.0e+5 2.5e+5 3.0e+5 3.5e+5 tds (mg/l) = 1.656 × cl (mg/l) – 3528 (mg/l) r²= 0.997 https://doi.org/10.34194/62417j08 http://www.geusbulletin.org/ schovsbo et al. 2025: geus bulletin 60. 8383. https://doi.org/10.34194/62417j08 7 of 17 geusbulletin.org measurements could only be made for the stenlille area (fig. 1) where the predicted pw in the gassum formation using eq. 3 and measured pw in the stenlille-5 well is within 4 bar. for the more saline and hence more dense brine in the bunter sandstone and skagerrak formations, the estimated hydrostatic pressures should be regarded as minimum pressure estimates. formation temperatures are calculated from fuchs et  al. (2019). for the stenlille area, we assume that the temperature model for stenlille-1 (fuchs et  al. 2019) applies to all stenlille wells. for gassum-1 and rødby-1, where no thermal gradients were reported by fuchs et al. (2019), we use: t (°c) = 0.027 (°c/m) · [depth in m below ground level] + 8°c (4) where the thermal gradient is 27°c/km and the surface temperature is 8°c as used in the deep geothermal evaluation of denmark (vosgerau et al. 2016). 3.2.4. brine viscosity brine (nacl) viscosity (μbrine), is calculated as a function of temperature and nacl concentration according to the correlation developed by phillips et al. (1981): μbrine = (1 + a·m + b·m2 + c·m3 + d·t· (1 – ekm)) · μw (5) where a = 0.0816, b = 0.0122, c = 0.000128, d = 0.000629, k = –0.7, t = temperature (°c), m = molal concentration (mol nacl/kg h2o), μw = viscosity of water (cp). the concentration of nacl is calculated from the tds content assuming a pure nacl solution. μw is estimated from tables provided by the national institute of standards and technology by the u.s. department of commence (linstrom & mallard 2023). francke & thorade (2010) compared the relationship given in eq. (5) to models by kestin et al. (1981) and mao & duan (2009) and demonstrated consistent values and a deviation below 0.9%. 3.2.5. co2 density and viscosity co2 density (δco2) and viscosity (μco2) at reservoir conditions are estimated from tables provided by the nist chemistry webbook. 3.2.6. partial least squares regression analysis the partial least squares (pls) multivariate regression method was employed to explore relationships between chloride concentrations and depth, in relation to the main stratigraphic tops and facies. the multivariate x data were prepared from depths in relation to the listed stratigraphic tops in table 1. thus, the depth of a brine sample was calculated relative to any given datum by determining its depth above (positive value) or below (negative value) the datum, resulting in a matrix termed as the spatial relationship matrix (srm; fig. 5). this approach facilitates the development of predictive models (see schovsbo et al. 2020). the pls regression enables direct modelling of correlations between the dependent cl variable (y) and the multivariate table 1 coefficients for calculation of b in eq. 6 for wells. depth reference is ground level. depth reference abbreviation coefficient depth below ground level dbt a1 0 depth below top chalk group dbt_chalk a2 6.1312 depth below top lower cretaceous dbt_lc a3 20.2594 depth below upper jurassic dbt_ujur a4 14.3422 depth below mid jurassic dbt_mjur a5 6.4713 depth below top triassic (gassum fm) dbt_tr a6 -14.1453 depth below top mid triassic (vinding fm) dbm_tr a7 -16.4221 table 2 brine properties within structures and selected samples. structure / well formation depth (tvdgs) cl tds cl pls tds pls pw t co2 solubility sinacl m mg/l mg/l mg/l mg/l bar oc g co2/l st-1 fjerritslev 1366 103 000 169 523 144 50 35.1 –0.81 stenlille structure gassum 1506 106 081 175 670 159 54 st-1 gassum 1506 108 000 177 894 159 54 33.9 –0.76 st-19 falster 2053 182 000 302 969 216 71 23.1 0.18 stenlille structure bunter sandstone 2413 181 249 300 148 254 82 st-19 bunter sandstone 2413 197 000 322 452 254 82 21.5 0.23 gassum structure frederikshavn 1095 99 437 164 667 115 38 gassum structure gassum 1545 115 206 190 782 163 50 ga-1 ørslev c. 2625* 180 770 287 986 277 79 24.4 –0.03 gassum structure skagerrak 2795 201 803 334 185 295 83 rødby structure bunter sandstone 1246 186 265 308 454 131 42 inez structure gassum 1650 107 416 177 881 179 50 lisa structure gassum 1669 147 772 244 710 180 47 *depth uncertain. tvdgs: total vertical depth below ground level / sea floor. tds: total dissolved solids. cl pls: chloride concentration estimated by the pls regression model. pw: hydrostatic pressure at reservoir depth. t: temperature at reservoir depth. co2 solubility: co2 in brine at reservoir conditions. sinacl: saturation state of brine–halite at reservoir conditions. https://doi.org/10.34194/62417j08 http://www.geusbulletin.org/ schovsbo et al. 2025: geus bulletin 60. 8383. https://doi.org/10.34194/62417j08 8 of 17 geusbulletin.org independent depth variables (x), based on the collinearity among x variables (esbensen & swarbrick 2018). the pls regression model was cross-validated by randomly splitting the data into two segments – an independent training set and a test data set – ensuring a realistic prediction performance validation (esbensen & geladi 2010; esbensen & swarbrick 2018). all data were auto-scaled, and the modelling was performed in the software package unscrambler® 10.5 by camo (esbensen & swarbrick 2018). 4. results 4.1. salinity prediction laier (2008) discussed methods for predicting water composition in reservoirs lacking direct analyses. one approach, originally presented by laier (1989a), involves using a simple model that correlates chloride concentration with present-day depth by plotting cl concentrations versus depth. figure 6a presents a similar depth plot for the full augmented database (supplementary table s1). the overall relationship has an r² of 0.65, reflecting that only 65% of the total data variance can be modelled by a simple correlation relationship (fig. 6a). as discussed by laier (1989a), the depth relationship in fig. 6a encompasses several sub-trends: 1. the tønder samples, representing partly halite-cemented bunter sandstone (laier & nielsen 1989; hjuler et al. 2019), plot with high and apparently invariant cl–depth relationships. 2. the erslev-4s data exhibit a steep cl–depth gradient, reflecting the intrusion of shallow salt diapirs into the ground water zone. 3. the bulk of the gassum and younger reservoirs plot with a relatively well-defined cl–depth relationship. laier (1989a) also noted the wide range of chloride levels in the skagerrak samples, from 110 000 ppm at margretheholm (2550 m depth) to 200 000 ppm at stenlille-19, despite both being at nearly the table 3 density and viscosity of brine and scco2. structure / well formation depth (tvdgs) δw surface δw δco2 δw – δco2 δwco2 μco2 μbrine μbrine / μco2 m kg/l kg/l kg/l kg/l kg/l cp cp st-1 fjerritslev 1366 1.12 1.10 1.11 stenlille structure gassum 1506 1.12 0.69 0.43 0.055 0.71 12.8 st-1 gassum 1506 1.12 1.11 1.11 st-19 falster 2053 1.21 1.18 1.18 stenlille structure bunter sandstone 2413 1.21 0.68 0.53 0.056 0.63 11.2 st-19 bunter sandstone 2413 1.22 1.20 1.20 gassum structure frederikshavn 1095 1.11 0.73 0.39 0.060 0.91 15.2 gassum structure gassum 1545 1.13 0.73 0.40 0.061 0.80 13.0 ga-1 ørslev c. 2625* 1.20 1.18 1.18 gassum structure skagerrak 2795 1.23 0.73 0.50 0.062 0.66 10.6 rødby structure bunter sandstone 1246 1.21 0.73 0.49 0.060 1.14 19.0 inez structure gassum 1650 1.12 0.75 0.37 0.064 0.76 11.8 lisa structure gassum 1669 1.17 0.78 0.39 0.068 0.89 13.1 *depth uncertain. tvdgs: total vertical depth below ground level / sea floor. δw surface: density of formation water at ambient conditions. δw: density of formation water at reservoir conditions. δco2: density of scco2 at reservoir conditions. δw – δco2: density difference between formation water and scco2. δwco2: density of co2 saturated formation water at reservoir conditions. μco2: viscosity of scco2 at reservoir conditions. μbrine: viscosity of brine (nacl) at reservoir conditions. fig. 5 principle of the spatial relationship matrix (srm). the transformation expresses the sample depth relative to a specific surface by calculating its depth above (positive value) or below (negative value) the datum. consequently, a sample in a well with ‘n’ identified surfaces will have n depths. s1: surface 1. s2: surface 2. s3: surface 3. s4: surface 4. d ep th u ni t c u ni t d u ni t b u ni t a st ra ti gr ap hy spatial relationship matrix surface 2 surface 3 surface 4 surface 1 elevation above s3 (+) elevation above s4 (+) depth below s2 (–) depth below s1 (–) sample https://doi.org/10.34194/62417j08 http://www.geusbulletin.org/ schovsbo et al. 2025: geus bulletin 60. 8383. https://doi.org/10.34194/62417j08 9 of 17 geusbulletin.org same  depth. these variations render the depth plot alone less effective for predicting chloride concentrations in unsampled reservoirs. consequently, laier (2008) proposed an alternative approach that combines analogue considerations between reservoirs of known and unknown compositions, applying adjustments for depth differences based on laier (1989a). using this method, the likely brine compositions for seven reservoirs at specific locations and depths were determined. 4.2. partial least squares regression model of salinity 4.2.1. modelling concepts to develop a mathematical model for the relationship between chloride concentrations and depth, we began by identifying the fundamental controls influencing salinity in formation brines. to determine if the variation is controlled by saturation-induced constraints, we calculated the saturation state of the formation water with respect to halite (brine–halite si) under reservoir conditions. we then plotted this against the depth below ground level for water samples with a full compositional analysis (fig. 6b). the analysis indicates that, except for the very deep (>2.5 km) reservoirs in aars-1a and farsø-1, all samples from the gassum and younger reservoirs are undersaturated with respect to halite. in contrast, all water samples from the bunter sandstone, skagerrak, ørslev and falster formations (except those from margretheholm), and from the zechstein group, are in equilibrium with halite (fig. 6b). to determine the geographical component, we observed that the salinity variation in the bunter sandstone and skagerrak formations differs between the north german basin, the danish basin and the øresund basin that forms a marginal part of the danish basin. within these basins, the tønder area in the north german basin has highest salinities, and the margretheholm and höllviksnäs areas in the øresund basin are the least saline (fig. 6a). the subsurface rock salt occurrence presented in fig. 3a, b, c was compared with sample locations for water chemistry (fig. 7). we observed that, for the bunter sandstone and skagerrak formations, the primary control on salinity levels appears to be stratigraphic proximity to rock salt in the zechstein group and the ørslev and oddesund formations as rock salt is present in the north german basin and danish basin but absent in the øresund basin. to account for this, we incorporated information on the presence or absence of rock salt into the pls regression model reflecting the proximity to, or absence of, interbedded rock salt layers within each basin. this was fig. 6 chloride concentrations versus present day depth below ground level (a) and brine-halite (nacl) saturation index (sinacl) at reservoir conditions (b). thermodynamic equilibrium for halite (–0.4 ≤ sinacl ≤ 0.4) is illustrated with a grey band. well abbreviations as fig. 1. broken lines in a indicate sub-trends discussed in the text. the dge-1 well tested crystalline basement between 3198 and 3702 m of depth. aarhus-02 (aar-02) tested the interval 2293–2411 m. cl (mg/l) = 36 665 mg/l + 49.7 mgl-1/m [depth below ground] r²= 0.65 erslev trend tønder trend gassum trend cl (mg/l) 0 100000 200000 300000 d ep th b el ow g ro un d le ve l ( m ) 0 250 500 750 1000 1250 1500 1750 2000 2250 2500 2750 3000 3250 3500 hö-1 d g e1 equilibriumunder saturated saturated brine-halite saturation index –1.5 –1.0 –0.5 0.0 0.5 1.0 aa-1a st-19 th-2 st-19 tø-4,-5 ma-1 ma-2 fa-1 aa-1a ga-1 gassum and younger reservoirs ørslev/falster/bunter/skagerrak/ zechstein linear regression aa-1a st-19 th-2 st-19 ør-1 tø-4,-5 ma-1 ma-2 er-4s fa-1 th-2 aa-1a ga-1 aar-02 a b https://doi.org/10.34194/62417j08 http://www.geusbulletin.org/ schovsbo et al. 2025: geus bulletin 60. 8383. https://doi.org/10.34194/62417j08 10 of 17 geusbulletin.org accomplished by introducing parameters labelled cngb and cdg in fig. 7, representing categorical contribution variables from the north german basin and the danish basin, respectively. to model the depth trend in the chloride concentrations, we employed an analysis of the stratigraphical data from the wells by calculating the depth of the brine samples to the given stratigraphical surfaces resulting in a srm of the brine sample depths. this approach follows that of schovsbo et al. (2020), who modelled depth trends in vitrinite reflectance data from north sea wells by incorporating stratigraphical data. the srm (see fig. 5) enables stratigraphical data to be modelled using multivariate statistical tools, such as pls regression. 4.2.2. partial least squares model results for preparation of the srm, stratigraphical information was obtained for danish wells from nielsen & japsen (1991) with updates from gregersen et  al. (2023), abramowitz et  al. (2024), keiding et  al. (2024) and for swedish wells from erlström & sivhed (2012) and erlström et al. (2018). the eg(s)-1 well, the shallow research well stevns-1 that only penetrated the upper part of the chalk group (stemmerik et al. 2006) and the haldager-1 well that terminated in the fjerritslev formation were not included due to limited stratigraphical coverage. in addition, the dge-1 well that tested the crystalline basement between 3198 and 3702 m in scania was excluded due to uncertainties concerning the depths of the water analyses. in the srm, we used the sample depth calculated in relation to the following surfaces: ground level, top chalk group, top lower cretaceous, top upper jurassic, top middle jurassic, top triassic (equivalent to the top gassum formation) and the top middle triassic (equivalent to top vinding formation or base gassum formation). to account for the chloride contribution from interbedded rock salt in the zechstein group and the ørslev and oddesund formations to the bunter sandstone, falster and skagerrak reservoirs, categorical contribution variables, cngb and cdb were assigned according to fig. 7. a cngb value of 1 indicates the presence of the well within the depositional areas of the zechstein group and the ørslev formation in the north german basin, while a cdb value of 1 signifies the presence of the well within the depositional area of rock salt from the zechstein group and/or the oddesund formation in the danish basin (fig. 7). wells outside these areas, such as margretheholm-1 and -2 and swedish wells in the øresund basin (erlström et al. 2018), were assigned a cdb value of 0. for gassum and younger reservoirs, both cngb and cdb are assigned a value of 0. the resulting pls regression model of the chloride concentrations is presented in fig. 8. the full input variable [srm, cngb, cdb] data are available in supplementary table s2. the proportions of total data variance modelled shown along each pls component [x%, y%] are shown in fig. 8b. from this diagram, a three-component pls model on the [srm, cngb, cdb] variable set predicts chloride levels with very satisfactory validation results, as seen in the prediction versus reference plot in fig. 8c (slope 0.95; r2 = 0.95). hence, the pls regression model exhibits a correlation between measured chloride levels and predicted chloride levels that is much stronger than could be established from using only the depth below surface for predicting the chloride concentrations (r2 = 0.65; fig. 6a). the resulting pls regression model is described by: gassum and younger reservoirs: clgassum (mg/l) = 35.04 mgl–1/m · [depth below ground level in m] + (56 259 – b) (mg/l) (6) where b is a well-specific constant calculated from: b = a1·dbt + a2·dbt_chalk + a3·dbt_lc + a4·dbt_ujur + a5·dbt_mjur + a6·dbt_tr + a7· dbm_tr (7) using coefficients for a1–a7 for the stratigraphical depths calculated in the well as presented in table 1. values for b for the analysed wells are presented in supplementary table s3. for the bunter sandstone, falster, ørslev and skagerrak reservoirs, an additional salinity component is calculated in addition to the chloride calculated from eq. (6). fig. 7 assignment of the categorial contribution variables for north german basin (cngb) and danish basin (cdb) for the pls regression analysis of zechstein, bunter sandstone, ørslev, falster and skagerrak brine samples. a value of 1 indicates the likely presence of interbedded rock salt and a value of 0 indicates the likely absence. assessed potential co2 structures (yellow) and wells are after fig. 1. 56°n 7°e 55°n 56°n 57°n 8°e 9°e 10°e 11°e 12°e 100 km röt salt röt salt risk (webgis) zechstein platform zechstein salt basin oddesund salt cngb = 1 cdb = 1 cngb and cdb= 0 bunter / skagerrak https://doi.org/10.34194/62417j08 http://www.geusbulletin.org/ schovsbo et al. 2025: geus bulletin 60. 8383. https://doi.org/10.34194/62417j08 11 of 17 geusbulletin.org clbunter/skagerrak (mg/l) = clgassum (mg/l) + 86 240 (mg/l) · cngb + 42 792 (mg/l) · cdb (8) where cngb and cdb are assigned according to fig. 7. a reference salinity calculator based on eqs (6), (7) and (8) is included in the supplementary files. 4.3. assessment of selected reservoirs in co2 structures further in the text we have evaluated the salinity (cl and tds), t, pw, δw, δco2, μco2, μbrine, halite saturation state and the co2 solubility in brine for selected reservoirs in five structures that are identified as suitable for co2 storage (gregersen et  al. 2023, 2025, this volume; abramowitz et  al. 2024; keiding et  al. 2024). in our selection, we focused on drilled structures to establish stratigraphical control for the modelling. however, this is not a technical hindrance, as a pseudo well could also have been used for the assessment (see bjerager et al. 2024 for an example). the depth profiles of the assessed chloride levels are depicted in fig. 9 and all evaluated parameters are presented in tables 2 and 3. 4.3.1. the stenlille structure the stenlille structure is situated in the central part of the island of sjælland in the danish basin (fig. 1). the structure is a four-way closure developed on top of a zechstein salt pillow and is currently both an active natural gas storage site and a potential co2 storage site, with the gassum and bunter sandstone formations identified as injection possibilities (gregersen et  al. 2023). twenty brine analyses are available from seven wells (fig. 1, supplementary table s1). we use the reservoir depths from the stenlille-1 and stenlille-19 wells, and the full composition analysis from these wells for the fjerritslev, gassum, falster and bunter sandstone formations as references (table 2). at stenlille, the pls-predicted chloride concentration for the gassum formation matches the available composition within uncertainty (pls 106 000 vs. measured 108 000 mg/l). however, for the bunter sandstone formation at 2413 m, the pls-estimated cl content is 16 000 ppm too low (181 000 vs. 197 000 mg/l; table 2). brine density at surface conditions estimated using eq. (4) is up to 0.03 kg/l lower than that estimated with phreeqc based on the measured brine composition at reservoir conditions (table 3). the shallowest water samples at stenlille, measured in the fjerritslev and gassum formations, are undersaturated with respect to halite with sinacl of –0.81 to –0.76, whereas the deeper water samples measured in the falster and bunter sandstone formations are in equilibrium with respect to halite (si = 0.18 to 0.23). due to increasing salinity with depth, co2 solubility also decreases with depth. thus, co2 solubility decreases from approximately 34 g co2/l in the gassum formation at 1506 m to 21.5 g co2/l in the bunter sandstone formation at 2413 m (table 2). supercritical co2 will have a density of about 0.68 kg/l in both the gassum and bunter sandstone formations. due to the higher tds in the latter formation, the water density here is highest, resulting in a relative density difference between the two phases of 0.43 kg/l in the gassum formation and 0.53 kg/l in the bunter sandstone formation. the viscosity (μbrine) is in the range 0.63–0.71 centipoise (cp), and the mobility ratio, calculated as the ratio between brine and co2 viscosities, ranges from 11.2 to 12.8 with the lowest ratio in the bunter sandstone formation due to its high temperature (table 3). the relatively high mobility ratios may increase the pls factors 0 1 2 3 4 5 6 7 yva ri an ce 0 20 40 60 80 100 measured cl (mg/l) 0.0 5.0e+4 1.0e+5 1.5e+5 2.0e+5 2.5e+5 pr ed ic te d c l ( m g/ l) 0.0 5.0e+4 1.0e+5 1.5e+5 2.0e+5 2.5e+5 b cr² = 0.95 y = 5679 + 0.95 x root mean square error: 7963 pls 1 (70%, 74%) -0.4 -0.2 0.0 0.2 0.4 pl s 2 (1 0% , 1 5% ) -0.6 -0.4 -0.2 0.0 0.2 0.4 0.6 0.8 a dbt dbt_chalk dbt_lc dbt_ujur dbt_mjur dbt_tr dbm_tr cdb cngb cl fig. 8 pls-regression model for chloride concentrations (red) and stratigraphic and geologic parameters [srm, cngb, cdb] variable set (black dots). (a) loading-weights plot (w1–w2); proportions of total data variance modelled shown along each pls component [x%, y%]. (b) modelled y-variance and (c) prediction versus reference plot. proportions of total data variance (pls factor 3) modelled shown along each pls component [x%, y%]. categorial contribution variables for north german basin (cngb) and danish basin (cdb). cl: chloride. other abbreviations after table 1. https://doi.org/10.34194/62417j08 http://www.geusbulletin.org/ schovsbo et al. 2025: geus bulletin 60. 8383. https://doi.org/10.34194/62417j08 12 of 17 geusbulletin.org risks of viscous fingering. the relative increase in formation water density at a co2-saturated state (δwco2) is 0.3% for the gassum formation and negligible for the bunter sandstone formation (table 3). 4.3.2. the gassum structure the gassum structure is situated in central jylland within the danish basin and is a four-way closure atop a zechstein salt pillow (fig. 1). the structure was drilled by the gassum-1 well and is targeted for co2 storage, with the frederikshavn, gassum (primary) and skagerrak formations as possible reservoirs (keiding et al. 2024). in the gassum-1 well, a brine sample was retrieved during drilling due to uncontrolled well flow, as reported in the final well report (danish american prospecting 1951, p. 324). due to the conditions under which the sample was collected, the exact depth remains uncertain but is believed to represent ‘sands logged at and below 8816 feet (2687 m)’. laier (2008) included this sample in his compilation, assigning it a depth of 2625  m and to the falster/skagerrak formation. we have retained this estimate while noting that the actual depth could be somewhat deeper and consider it to represent an ørslev formation brine following keiding et al. (2024; tables 2 and 3). applying the pls regression model with the gassum-1 stratigraphy as input, it is estimated that at the depth of the frederikshavn formation, situated at approximately 1095 m, the brine has a chloride concentration of 99 400 mg/l (equivalent to 164 700 mg/l tds). within the primary reservoir of the gassum formation, located at roughly 1550 m below ground level, the estimated salinity reaches approximately 115 200 mg/l cl (190 700 mg/l tds). meanwhile, at the depth of the skagerrak sandstone reservoir, around 2800 m, salinity is estimated to be approximately 201 800 mg/l cl (334 100 mg/l tds) with a density of 1.23 kg/l. the influx at 2625 m in the ørslev formation is modelled to have a cl level of 195 000 mg/l, which is 15 000 mg/l higher than measured values of 180 770 mg/l cl, 287 986 mg/l tds. given the uncertainties with the depth of the influx and modelling, this seems reasonable. halite saturations and co2 solubility cannot be modelled for the three reservoir levels. from fig. 9, we assume that the brine compositions are undersaturated with respect to halite in the frederikshavn and gassum formations and in equilibrium with respect to halite in the skagerrak formation. the influx at 2625 m in the ørslev formation is modelled to be in equilibrium with halite and to have a co2 solubility of 24 g co2/l. at this level, the increase in δw upon saturation with co2 is negligible (table 3). values of δco2 in all reservoirs are around 0.73 kg/l, which means that the density contrast between co2 and brine is highest (0.50 kg/l) in the skagerrak formation and lowest in the frederikshavn formation (0.40 kg/l). the μbrine ranges from 0.66 to 0.91 cp and the mobility ratio (μbrine/μco2) ranges between 10.6 and 15.2 with the lowest value in the skagerrak formation. 4.3.3. the rødby structure the rødby structure is situated in the western part of the island of falster in the north german basin (fig. 1). this structure is a four-way closure developed on a zechstein salt pillow and is targeted for co2 storage within the bunter sandstone formation, as the gassum formation is too shallow to be suitable for co2 storage (abramowitz et al. 2024). we use the rødby-1 well data as the basis for our analysis (table 2) and since no brine composition analyses are available, we have used the salinity (cl and tds) modelled from the pls regression model (eq. 6) as the basis for the formation water composition. accordingly, we estimate the bunter sandstone brine cl concentration at a depth of 1246 m to be 186 000 mg/l, the tds to be 308 000 mg/l and the δw at reservoir conditions to be 1.21 kg/l (table 2). no saturation simulation of halite or co2 solubility could be conducted. however, comparing the depth relationship of chloride and the sinacl, we assume that the brine in rødby is in equilibrium with halite (fig. 9) and that the solubility of co2 in brine is slightly less than that in the bunter sandstone formation in stenlille due to lower hydrostatic pressures (pw is 131 bar in rødby versus 254 bar in stenlille) and temperatures (42°c in rødby and 82°c in stenlille) in the bunter sandstone reservoir in rødby compared to stenlille (table 2). the co2 in rødby is assumed to have a δco2 value at reservoir conditions of 0.73 kg/l, and the density difference between scco2 and brine is modelled to be 0.49 kg/l. the μbrine is modelled to 1.14 cp due to the high salinity and relatively low temperature (42°c). combined with μco2, this gives a mobility ratio of 19.0. 4.3.4. the lisa structure the lisa structure is situated offshore, approximately 60 km west of northern jylland, in the north sea area within the danish basin (fig. 1). the structure was drilled by the j-1x well and is a four-way closure upon an upper triassic salt pillow, targeted for co2 storage with the gassum formation as the primary reservoir (gregersen et al. 2025, this volume; see also fyhn et al. 2024). no formation water analysis is available for the structure. using the j-1x stratigraphical data as input, we evaluate the salinity in the gassum formation at 1650 m below ground level to be 147 700 mg/l, corresponding to 244 700 mg/l tds, with a δw of 1.12 kg/l (table  2). https://doi.org/10.34194/62417j08 http://www.geusbulletin.org/ schovsbo et al. 2025: geus bulletin 60. 8383. https://doi.org/10.34194/62417j08 13 of 17 geusbulletin.org since no phreeqc simulation can be made, we use fig. 9 to assume that the brine is undersaturated with respect to halite. the co2 solubility is expected to be between 30 and 35 g co2/l, as seen in the fjerritslev and gassum formations in the stenlille structure (table 2). at this pressure (pw) and temperature (180 bar, 47°c), δco2 will be 0.75 kg/l, and the density difference to the brine will be around 0.40 kg/l. the μbrine is modelled to be 0.89 cp, which combined with μco2 has a mobility ratio of 13.1 (table 3). 4.3.5. the inez structure the inez structure is situated offshore, approximately 80 km west of jylland, in the north sea within the danish basin (fig. 1). the structure was drilled by the inez-1 well and is a four-way closure upon a zechstein salt structure, targeted for co2 storage within the gassum formation (gregersen et al. 2025, this volume). no formation water analyses are available for the structure. using inez-1 as stratigraphical input data, we evaluate the salinity in the brine at the level of the gassum formation at 1650 m below ground level to be 107 400 mg/l, corresponding to 177 800 mg/l tds (table 2). based on fig. 9, we assume that the brine is undersaturated with respect to halite. the co2 solubility is expected to be in the range of 30–35 g co2/l, similar to the stenlille structure (table 2). co2 in the structure will have a density of 0.77 kg/l, and the density difference to the brine will be around 0.37 kg/l. the μbrine is modelled to 0.76 cp, which combined with μco2 has a mobility ratio of 11.8 (table 3). 5. discussion 5.1. origins and dynamics of brine sources the increase in salinity with depth shown in fig. 6a is typical for basins with deep-bedded salt or brines derived from the subaerial evaporation of seawater (kharaka & hanor 2003) and is thus unsurprising given the nature of the danish area (fig. 3). laier (1989a) highlighted the existence of diverse brine sources and employed comprehensive chemical analyses to confirm their presence. he observed that the salinity and composition of the bunter sandstone – skagerrak formation brine were significantly influenced by residual brines from massive salt precipitation during the late permian (zechstein) period. these residuals were hypothesised to have migrated or facilitated ion exchange through diffusion, markedly impacting the salinity and brine compositions. furthermore, laier noted that the trend of increasing salinity with depth could stem from fluid mixing or hyperfiltration during sediment burial. he argued that the clays deposited during the deposition of the vinding formation may have impeded the mixing of pre-upper triassic waters with younger formation waters, which exhibit distinctly lower br/cl ratios. the younger waters (gassum formation and younger) probably evolved from seawater through hyperfiltration. holmslykke et al. (2019) categorised the brine compositions based on multivariate data analysis. they identified three main types with further subdivisions: type 1a (low tds, sulphate-rich), type 1b (low tds, sulphate-depleted), type 2a (ca–mg rich, low k), type 2b (ca–mg rich, high k) and type 3 (na–cl rich). in the water typing, chloride was seen as a conservative ion. the main variation in composition was due to local diagenesis and structural dependencies. accordingly, no specific water type could be assigned to a reservoir age. instead, subtypes dominated in specific sites, such as a ca–mg-low k-rich brine (type 2a) in the margretheholm area, a na–cl-rich type (type 3) in the tønder area and a sulphate-rich type (1a) in the sønderborg area. the most widespread type present at >2 km depth was a ca–mg–k-rich type that occurs in the haldager sand, gassum, falster and bunter sandstone reservoirs. our pls regression model includes primary control elements on the brine salinity, namely the proximity of evaporitic deposits to the bunter sandstone and skagerrak reservoirs (fig. 7). analysis of the geological surfaces using the srm allows us to compensate for some aspects of the geological development, such as fig. 9 measured chloride concentrations and pls-modelled chloride concentrations for measurements and reservoir levels in structures versus depth below ground level. field of brine–halite equilibrium (grey) is based on phreeqc simulations (fig. 6b). for well abbreviations, see fig. 1. cl (mg/l) 0 100000 200000 300000 d ep th b el ow g ro un d le ve l ( m ) 0 250 500 750 1000 1250 1500 1750 2000 2250 2500 2750 3000 3250 3500 aa-1a rødby structure inez structure lisa structure stenlille structure gassum structure brine-halite equilibirum pls model modelled measurements ga-1 st-19 st-19 ør-1 tø-4,-5 ma-1 ma-2 hö-1 gassum and younger reservoirs ørslev/falster/bunter/ skagerrak/zechstein measurements https://doi.org/10.34194/62417j08 http://www.geusbulletin.org/ schovsbo et al. 2025: geus bulletin 60. 8383. https://doi.org/10.34194/62417j08 14 of 17 geusbulletin.org variation in the thickness of the geological layers, which has resulted in a much-improved prediction of the depth trend of chloride concentrations (figs. 8 and 9). however, the variation in water types has shown that these cannot be reduced to just a matter of salinity as the water is highly evolved through water–rock interactions as shown by holmslykke et al. (2019). this implies that despite a good understanding of salinities, many aspects of the brine cannot be simulated through programs such as phreeqc unless a full water composition is known. only with complete compositions can the impact of co2 injection into reservoir brines be fully assessed, highlighting the need for detailed, localised geological and chemical analyses to understand the complex interplay of geological, chemical and physical processes that shape the hydrological characteristics of sedimentary basins. 5.2. implications for co2 storage the physicochemical characteristics of the co2–brine– rock system are of primary importance for operating a site to ensure safe and sustainable operations. our analysis of the five co2 storage sites in denmark (stenlille, gassum, rødby, lisa and inez), encompassing eight reservoirs (one in the frederikshavn formation, four in the gassum formation and three in the bunter sandstone and skagerrak formations) has shown a large range in values. the water characteristics vary substantially due to differences in temperature (38–83°c), pressure (115– 295 bar) and total salinity (165 000–334 000 mg/l tds), without even including the reservoir characteristics of the frederikshavn, gassum and bunter sandstone and skagerrak formations that are known to vary with location and depth (see pre-drilling assessment of porosity and permeability by kristensen et al. 2016; weibel et al. 2017; hjuler et  al. 2019; olivarius et  al. 2019). some implications for co2 storage are discussed further in the text, although conceptual fluid-flow modelling is recommended to further assess brine–rock–co2 interactions. 5.2.1. salt precipitation one of the challenges accompanying injection of dry scco2 is the risk of salt precipitation as water evaporates into the unsaturated co2. research indicates that both porosity and permeability reductions are to be expected, although estimates of the impact vary considerably (see edem et al. 2022; cui et al. 2023). this will affect both storage capacity and injectivity. the bunter sandstone and skagerrak reservoirs, as well as deep gassum reservoirs (>2.5 km), already in equilibrium with halite (fig. 6b), have a high risk in this respect, and injection schemes must account for this for successful co2 injection operations. however, near-wellbore salt precipitation can occur even in low-salinity reservoirs (cui et  al. 2023), which also presents a risk factor in the danish central graben reservoirs (narayanan et al. 2023) that have significantly lower salinities than the eastern danish reservoirs studied here (see schovsbo et al. 2016). therefore, the risk of salt precipitation should be viewed as an operational challenge; managing injection flow rates, choosing between continuous versus discontinuous flow and pumping precipitation inhibitors are key strategies to mitigate the risk (ringrose 2020; cui et al. 2023). 5.2.2. co2 solubility in brine the solubility of co2 in formation waters depends on temperature, pressure and water composition. as temperature and pressure increase, co2 solubility also increases, whereas high salinity lowers solubility. the dissolution of co2 in formation water is controlled by these factors and continues over millennia, contributing to the long-term storage fate of the injected co2 along with its mineralisation from reactions with the rock itself. high dissolution rates can significantly enhance the overall efficiency of co2 sequestration (holloway 2005; deng et al. 2018). our phreeqc modelling suggests a range in co2 dissolution from 21.5 to 33.9 g co2/l (table 2) – although not all structures were assessed due to the lack of a full composition analysis of the brine. these estimates can serve as initial points in the analysis of storage efficiency. another outcome of the modelling is the observation that due to the generally high densities of the brine, density increases of the water due to co2 saturation are low to negligible. in reservoirs with low salinity, the density difference between saturated and unsaturated brine could be 2–3%, sufficient to result in gravitational convection within the reservoir that would enhance the dissolution of co2 (kumar et al. 2020). in the case of the low-density contrast modelled here, such gravitational convection would be less pronounced. however, density-driven convection also depends on other factors such as reservoir height and permeability anisotropy (islam et al. 2016). for example, a thick, homogeneous sand layer would promote the development of optimal gravitational convection (frykman & wessel-berg 2014; jiang et al. 2019; wang et al. 2021). in the studied structures, we anticipate that gravitational convection is most likely to develop in low tds structures within the thickest reservoir sands. 5.2.3. viscosity and density brine viscosity (μbrine) is a key parameter influencing the ease with which fluids move through the porous media of a reservoir. higher viscosity indicates a more resistant fluid, while lower viscosity denotes a fluid that flows more readily. in the context of co2 storage, μbrine affects the injectivity and mobility of co2, determining how https://doi.org/10.34194/62417j08 http://www.geusbulletin.org/ schovsbo et al. 2025: geus bulletin 60. 8383. https://doi.org/10.34194/62417j08 15 of 17 geusbulletin.org easily co2 can be injected and how it displaces the resident brine within the reservoir. this, in turn, influences the displacement efficiency and pressure distribution during and after co2 injection. brine viscosity is highly temperature-dependent, decreasing at higher temperatures. additionally, increasing salinity leads to higher viscosity, adding complexity to reservoir management. our findings (table 3) show that brine viscosity ranges between 0.63 and 1.14 cp, with the ratio of μbrine to μco2 ranging from 11.8 to 19.0. while a comprehensive evaluation of viscosity effects requires consideration of relative permeabilities and capillary forces (berg & ott 2013), which is beyond the scope of this paper, these observations provide essential insights into reservoir behaviour. significant viscosity differences between brine and scco2 can lead to viscous fingering, impeding the effective distribution of co2 and affecting long-term storage integrity (kumar et al. 2020; ringrose 2020). the density contrast between brine and scco2 is another crucial factor influencing co2 storage. this contrast determines the buoyancy forces that drive the upward migration of co2 within the reservoir (tang et al. 2019). since scco2 is less dense than formation water, it tends to rise toward the top of the reservoir and potentially accumulate beneath the caprock. if it is not properly managed, this buoyant migration can increase the risk of leakage due to excessive gas column heights. furthermore, the density contrast affects the effective storage capacity of the reservoir. a higher density contrast means that co2 occupies a larger volume for the same mass, potentially reducing overall storage capacity. conversely, a lower density contrast allows co2 to be more densely packed, enhancing storage efficiency. our results indicate that the density contrast between scco2 and formation water is greater in the bunter sandstone and skagerrak formations compared to the gassum formation. this suggests a higher risk of co2 upward migration toward the caprock in the bunter sandstone and skagerrak formations, potentially leading to reduced storage capacity in these reservoirs. 6. conclusions the use of existing databases and the pls regression model for predicting salinity as a function of a set of differential depth distances to influencing evaporitic deposits has highlighted significant variations in water chemistry, which, when combined with temperature and hydrostatic pressure, result in large variations in key physical properties of formation brine, including its density, viscosity and co2 solubility. the analysis shows that for the studied structures, the lower to mid-triassic bunter sandstone and skagerrak formations and the deeply buried (>2.5 km) gassum formation have chloride levels ranging from 182 000 to 201 000 mg/l (303 000 to 334 000 mg/l tds). the bunter sandstone and skagerrak formations are significantly influenced by rock salt from the zechstein group and the lower to mid-triassic (falster and oddesund formations), whereas younger units such as the gassum and haldager sand formations and various cretaceous units, display salinity variations primarily driven by depth and localised geological conditions. formation water in all bunter sandstone – skagerrak reservoirs, with the exception of margretheholm, is in equilibrium with halite. consequently, the risk of salt precipitation due to reservoir evaporation in dry co2 is highest in these reservoirs. the differences between brine and scco2 densities and viscosities are significant, affecting the displacement efficiency of brine by scco2. the co2 dissolution modelled for five reservoirs ranges between 21.5 and 33.9 g co2/l, with a 37% decrease from high to low values due to chemistry and p and t differences. mobility ratios (μbrine/μco2) range between 10.6 and 19.0. complete water compositions are necessary to assess the full impact of co2 injection into reservoirs. this emphasises the need for detailed, localised geological and chemical analyses to fully understand the interplay of geological, chemical and physical processes that shape the hydrological characteristics of sedimentary basins. acknowledgments this study is part of the ccs2022–2024 project for maturation of selected structures to co2 storage sites. innargi is thanked for sharing data from the aarhus-02 well. the authors thank niels springer for fruitful discussions on pressure models. comments and suggestions from the reviewers philip ringrose and kim esbensen and the editor jon ineson are highly appreciated and significantly improved the quality of the paper. additional information funding statement this study is funded from the geus ccs2020–2024 project for maturation of selected structures to co2 storage site. author contributions nhs: conceptualization, data curation, formal analysis, investigation, methodology, project administration, resources, software, validation, visualization, writing – original draft writing – review & editing. hdh: conceptualization, data curation, formal analysis, investigation, methodology, software, validation, visualization, writing – original draft writing – review & editing. am: investigation, visualization, writing – review & editing. cmn: investigation, methodology, software, validation, writing – review & editing. competing interests no competing interests declared. additional files one excel spreadsheet containing supplementary table  s1 (geochemical database), table s2 (input file for the geochemical modelling) and table s3 (output files for geochemical modelling) is available at 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https://doi.org/10.34194/geusb.v35.4924 https://doi.org/10.37570/bgsd-2020-68-09 https://doi.org/10.34194/geusb.v10.4880 https://doi.org/10.1021/acsomega.9b00627 https://doi.org/10.1021/acsomega.9b00627 https://doi.org/10.22008/fk2/f9dwmb/imcmd5 https://doi.org/10.22008/fk2/f9dwmb/imcmd5 https://doi.org/10.34194/geusb.v35.4633 https://doi.org/10.34194/geusb.v35.4633 https://doi.org/10.1029/2020wr028132 https://doi.org/10.1029/2020wr028132 https://doi.org/10.1016/j.sedgeo.2004.05.004 https://doi.org/10.1016/j.geothermics.2016.09.003 https://doi.org/10.1016/j.geothermics.2016.09.003 https://doi.org/10.3390/geosciences14060146 assessment of formation brine salinity, pressure and temperature in selected structures in eastern denmark and implications for co2 storage 1. introduction 2. geological setting 3. data and methods 3.1. geochemical database 3.2. methods 3.2.1. estimation of total dissolved solids 3.2.2. brine density, halite saturation and co2 solubility 3.2.3. hydrostatic pressure and temperature 3.2.4. brine viscosity 3.2.5. co2 density and viscosity 3.2.6. partial least squares regression analysis 4. results 4.1. salinity prediction 4.2. partial least squares regression model of salinity 4.2.1. modelling concepts 4.2.2. partial least squares model results 4.3. assessment of selected reservoirs in co2 structures 4.3.1. the stenlille structure 4.3.2. the gassum structure 4.3.3. the rødby structure 4.3.4. the lisa structure 4.3.5. the inez structure 5. discussion 5.1. origins and dynamics of brine sources 5.2. implications for co2 storage 5.2.1. salt precipitation 5.2.2. co2 solubility in brine 5.2.3. viscosity and density 6. conclusions acknowledgments additional information funding statement author contributions competing interests additional files references figures fig. 1 location of wells and potential co2 storage sites (structures) mentioned in the text. well name abbreviations: er-4s: erslev-4s. fa-1: farsø-1. ga-1: gassum-1. ha-1: haldager-1. hö-2: höllviken-2. hö-1: höllviksnäs-1. in-1: inez-1. fcc-1: malmö/fcc-1. ma-1: margretheholm-1. ma-2: margretheholm-2. rø-1: rødby-1. st-1–6: stenlille-1, stenlille-2, stenlille-3 stenlille-4 stenlille-5 and stenlille-6. st-19: stenlille-19. stv-1: stevns-1. sø-1a: sønderborg-1a. sø-2: sønderborg-2. th-2: thisted-2. th-3: thisted-3. tø-1: tønder-1. tø-4: tønder-4. tø-5: tønder-5. ør-1: ørslev-1. aa-1a: aars-1a. aar-02: aarhus-02. fig. 2 mesozoic stratigraphy of the studied area. reproduced from gregersen et al. (2025, this volume). fig. 3 maps showing extension of rock salt in (a) the zechstein, (b) the röt (ørslev formation), and (c) the oddesund formation and (d) the depth to the pre zechstein surface. a is modified after geluk (2000), b after kovalevych et al. (2002) and vosgerau et al. 2016 (webgis röt salt risk map that highlights a high-risk area for rock salt presence), c after bertelsen (1980) and d after vejbæk & britze (1994). fig. 4 empirical relationships between chloride and tds concentrations established for mesozoic reservoirs. based on data in laier (2008) and holmslykke et al. (2019). fig. 5 principle of the spatial relationship matrix (srm). the transfor¬mation expresses the sample depth relative to a specific surface by calculating its depth above (positive value) or below (negative value) the datum. consequently, a sample in a well with ‘n’ identified sur¬faces will have n depths. s1: surface 1. s2: surface 2. s3: surface 3. s4: surface 4. fig. 6 chloride concentrations versus present day depth below ground level (a) and brine-halite (nacl) saturation index (sinacl) at reservoir conditions (b). thermodynamic equilibrium for halite (–0.4 ≤ sinacl ≤ 0.4) is illustrated with a grey band. well abbreviations as fig. 1. broken lines in a indicate sub-trends discussed in the text. the dge-1 well tested crystalline basement between 3198 and 3702 m of depth. aarhus-02 (aar-02) tested the interval 2293–2411 m. fig. 7 assignment of the categorial contribution variables for north ger¬man basin (cngb) and danish basin (cdb) for the pls regression analysis of zechstein, bunter sandstone, ørslev, falster and skagerrak brine samples. a value of 1 indicates the likely presence of interbedded rock salt and a value of 0 indicates the likely absence. assessed potential co2 structures (yellow) and wells are after fig. 1. fig. 8 pls-regression model for chloride concentrations (red) and strati¬graphic and geologic parameters [srm, cngb, cdb] variable set (black dots). (a) loading-weights plot (w1–w2); proportions of total data vari¬ance modelled shown along each pls component [x%, y%]. (b) modelled y-variance and (c) prediction versus reference plot. proportions of total data variance (pls factor 3) modelled shown along each pls component [x%, y%]. categorial contribution variables for north german basin (cngb) and danish basin (cdb). cl: chloride. other abbreviations after table 1. fig. 9 measured chloride concentrations and pls-modelled chloride con¬centrations for measurements and reservoir levels in structures versus depth below ground level. field of brine–halite equilibrium (grey) is based on phreeqc simulations (fig. 6b). for well abbreviations, see fig. 1. tables table 1 coefficients for calculation of b in eq. 6 for wells. depth reference is ground level. table 2 brine properties within structures and selected samples. table 3 density and viscosity of brine and scco2. japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 1 of 43 thermo-tectonic development of the wandel sea basin, north greenland peter japsen*1 , paul f. green2 , james a. chalmers1 1geological survey of denmark and greenland (geus), copenhagen, denmark, 2geotrack international, brunswick west, victoria, australia abstract the carboniferous–palaeogene wandel sea basin of eastern north greenland (north of 80°n, east of 40°w) is an important piece in the puzzle of arctic geology. it is particularly important for understanding how the paleocene–eocene convergence between greenland, the canadian arctic and svalbard relates to the compressional tectonics in the high arctic, collectively known as the eurekan orogeny. in this study, we present apatite fission-track analysis (afta) data and review published vitrinite reflectance data combined with observations from the stratigraphic record to place firmer constraints on the timing of key tectonic events. this research study reveals a long history of episodic burial and exhumation since the collapse of the palaeozoic fold belts in greenland. our results define pre-cenozoic exhumation episodes in early permian, late triassic, late jurassic and mid-cretaceous times, each involving the removal of kilometre-scale sedimentary covers. mid-paleocene exhumation defines the timing of compression along the major fault zones during the first stage of the eurekan orogeny, after the onset of sea-floor spreading west of greenland. regional exhumation that began at the end of the eocene led to the removal of most of a kilometre-thick cover that had accumulated during eocene subsidence and involved a major reverse movement along the harder fjord fault zone, northern peary land. these events took place after the end of sea-floor spreading west of greenland, and thus, represent post-eurekan tectonics. mid–late miocene exhumation is most likely a consequence of uplift and incision across most of the wandel sea basin study area. the preserved sedimentary sequences of the wandel sea basin represent remnants of thicker strata that likely extended substantially beyond the present-day outline of the basin. we find that the present-day outline of the basin with scattered sedimentary outliers is primarily the result of fault inversion during eurekan compression followed by deposition and removal of a kilometre-thick overburden. 1 introduction the carboniferous–palaeogene wandel sea basin in eastern north greenland is an important piece in the larger puzzle of arctic geology (figs. 1, 2). the basin records sedimentation along the western margin of svalbard and the barents sea, on the north-western flank of the danmarkshavn basin (offshore north-east greenland) and south of the cenozoic sea-floor spreading in the arctic ocean. several aspects of the development of the wandel sea basin remain controversial, in particular, the timing of a major, compressional tectonic event that affected the basin in post-cretaceous times is much debated. did it happen during a comparatively brief period of time around the *correspondence: pj@geus.dk received: 15 may 2019 accepted: 25 june 2020 published: 26 april 2021 keywords: eurekan orogeny, uplift, exhumation, svalbard, sverdrup basin abbreviations: afta: apatite fission-track analysis a.s.l.: above sea level cai: colour alteration index ctb: central tertiary basin egfz: east greenland fracture zone halip: high arctic large igneous province hffz: harder fjord fault zone kctz: kap cannon thrust zone tlfz: trolle land fault zone vr: vitrinite reflectance geus bulletin is an open access, peerreviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: catherine jex (geus, denmark) reviewed by: andy carter (birkbeck, university of london, uk), eckart håkansson (the university of western australia, australia) funding: see page 41 competing interests: 41 additional files: see page 41 research article https://doi.org/10.34194/geusb.v45.5298 https://orcid.org/0000-0003-1689-7820 https://orcid.org/0000-0001-5478-534x https://orcid.org/0000-0001-6651-2277 mailto:pj@geus.dk japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 2 of 43 www.geusbul let in.org 80°n 70°n 60°n 30°w 30°e 60 °w 120°w 90°w 0° 0 250 500 km 62–55 ma sea-floor spreading 55–35 ma 35 ma to present spreading centres transform fault arctic ocean lr wsb atlantic ocean iceland svalbard greenland labrador sea baffin bay canadian arctic archipelago fig. 1 cenozoic sea-floor spreading in the northern north atlantic area. sea-floor spreading between north america and greenland (62–55 ma) caused the first stage of the eurekan orogeny. during this time, greenland moved together with europe. spreading between europe and greenland (55–35 ma) resulted in the second stage of the eurekan orogeny. during this time, north america, greenland and europe moved as separate plates. seafloor spreading ceased west of greenland (at c. 35 ma) after which greenland moved together with north america and the eurekan orogeny ended. the danmarkshavn basin is located off northeast greenland between 76 and 82°n. lr: lomonosov ridge. wsb: wandel sea basin. modified from brozena et al. (2003), oakey & chalmers (2012) and gaina et al. (2017). cretaceous–palaeogene boundary (håkansson & pedersen 1982, 2001, 2015), during the eocene (von gosen & piepjohn 2003), during the late paleocene–early eocene (døssing et al. 2010; svennevig et al. 2016) or during the late cretaceous–paleocene (guarnieri 2015)? insights into the development of the wandel sea basin are thus important for providing a better understanding of the relationship between (1) the mid-paleocene to eocene movement of greenland relative to both north america (oakey & chalmers 2012) and eurasia (gaina et al. 2017) and (2) the palaeogene basin formation and compressional tectonics in the wandel sea basin (the kronprins christian land orogeny; håkansson & pedersen 2001, 2015) and in the neighbouring carboniferous–palaeogene basins in canada (the eurekan orogeny; okulitch & trettin 1991; embry & beauchamp 2019) and on svalbard (the west spitsbergen fold belt; steel et al. 1985; dallmann et al. 1993; dallmann 2015; jones et al. 2017). it is widely accepted that these tectonic events, which we refer to collectively as the eurekan orogeny, record the palaeogene convergence between greenland, the canadian arctic and svalbard (de paor et al. 1989; ricketts 1994; harrison 2008; oakey & chalmers 2012; dallmann 2015; gion et al. 2017; jones et al. 2017; embry & beauchamp 2019). others restrict the term eurekan orogeny to the time span in the eocene during which greenland was part of a separate plate that was able to move independently with respect to both north america and europe (e.g. piepjohn et al. 2016; tessensohn & piepjohn 2000). this definition, however, fails to include the deformations caused by the movement of greenland during the late paleocene. the presence of many isolated occurrences of sedimentary strata has been suggested to represent at least 20 late permian to mesozoic pull-apart basins formed in response to episodes of combined transtension and transpression along the plate boundary between north greenland and svalbard (håkansson & pedersen 2015). however, the possibility that these sub-basins may have been connected by substantial sedimentary covers that have since been removed by erosion has attracted little attention with some exceptions (stemmerik et al. 1998; pedersen et al. 2018). another unresolved issue is an enigmatic event that affected the north coast of kronprins christian land, in which upper cretaceous rocks reached extremely high vitrinite reflectance (vr) levels of 7–10% (håkansson et al. 1994). https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 3 of 43 www.geusbul let in.org in this study, we present new apatite fission-track analysis (afta®) data from northern and eastern greenland, which are focussed on the wandel sea basin. these data shed light on the thermal and tectonic history of the region, both prior to and after the compressional tectonics that affected the wandel sea basin in palaeogene times. we combine thermal history interpretations from afta data in 52 outcrop samples (of which 46 are from north of 79°n) with published vr data (håkansson et al. 1994; paech & estrada 2018; pedersen et al. 2018) to provide insights into the tectonic development of the region. afta data define the timing and magnitude of key palaeothermal episodes – times at which rocks cooled from higher temperatures compared with present-day values – while vr data provide independent determination of maximum post-depositional palaeotemperatures. we relate the thermal history solutions from afta and vr data to former depths of burial and corresponding amounts of exhumation and produce a synthesis of late palaeozoic to cenozoic thermo-tectonic events in the region. we integrate these results with observations from the geological record (håkansson & pedersen 2015; svennevig et al. 2016, 2017, 2018; alsen et al. 2018; hovikoski et al. 2018; paech & estrada 2018; pedersen et al. 2018; piasecki et al. 2018; bjerager et al. 2019). finally, we compare the results of this study with a companion study of north-east greenland (japsen et al. in press) and set the results in a regional context. 2 geological setting the wandel sea basin occupies a broad, partially submerged or ice-covered depression centred around the wandel sea (north of 80°n, east of 40°w; figs. 3 and 4; dawes & soper 1973; dawes 1976; håkansson & pedersen 1982, 2001, 2015; håkansson & stemmerik 1989; stemmerik et al. 1998). the basin is located between the caledonian fold belt along the east coast of greenland and the ellesmerian fold belt along the north coast (fig. 1; higgins et al. 2000). the main deformations took place during the silurian and the early carboniferous (henriksen et al. 2009). the wandel sea basin is the northernmost of a series of fault-bounded late palaeozoic – early palaeogene basins exposed along the eastern and northern margin of greenland (stemmerik et al. 1998). the basin contains sediments with a cumulative thickness of >15 km, resting unconformably on precambrian to silurian rocks (fig. 5; surlyk 1991; håkansson & pedersen 2015). upper carboniferous – triassic sediments accumulated in a system of grabens and half-grabens referred to as the north greenland – svalbard rift basin (håkansson & stemmerik 1989). the mesozoic basin evolution has been difficult to assess due to problems associated with correlating scattered outcrops and tectonic overprinting. however, recent studies have made such correlations possible. hovikoski et al. (2018) provided a jurassic–cretaceous fig. 2 outline of the high arctic adjacent to north greenland. extent of kronprins christian land orogeny according to håkansson & pedersen (2001) and this study. lr: lomonosov ridge. detail of fig. 1 with additions from von gosen & piepjohn (2003), oakey & chalmers (2012) and piepjohn et al. (2016). ellesmere island greenland greenland sea peary land kronprins christian land svalbard barents sea strike-slip fault thrust zone west spitsbergen fold belt kronprins christian land orogeny 500 km0 eurekan orogeny fault zone tlfzhffz kctz 80°n 30°w 30°e 60 °w 90°w 0° arctic ocean fram strait fig. 3 lr https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 4 of 43 www.geusbul let in.org lithostratigraphy for kilen (kronprins christian land), and bjerager er al. (2019) presented a triassic (induan– norian) lithostratigraphy for the wandel sea basin. håkansson & stemmerik (1989) suggested that the upper jurassic to cretaceous sediments accumulated in increasingly smaller sub-basins formed by strike-slip tectonics and transtension in what they called the ‘wandel hav strike-slip mobile belt’. according to håkansson & pedersen (1982, 2001, 2015), the ‘wandel hav strikeslip mobile belt’ is a north-west to south-eaststriking wrench-fault deformation zone that was instrumental in all basin-forming events from the late permian onwards and includes the following three, major fault systems (fig. 4): 1. the south-west to north-east-trending kap cannon thrust zone (kctz) is exposed over c. 75 km in northernmost peary land. ductile thrusting along the kctz took place between 49 and 47 ma, estimated from the ar-ar dating of volcanic rocks of the campanian to the earliest paleocene kap washington group (tegner et al. 2011; håkansson & pedersen 2015). 2. the east–west-trending harder fjord fault zone (hffz) is more than 300 km long and traverses peary k r o n p r i n s c h r i s t i a n l a n d kap washington kap cannon kap rigsdagen pr. ingeborg halvø pr. margrethe ø pr. dagmar ø pr. thyra ø holm land amdrup land herluf trolle land depotbugt hyde fjord independence fjord frigg fjordharder fjord nakkehoved station nord hovgaard ø kilen w a n d e l d a l p e a r y l a n d w a n d e l s e a 35°w 30°w 25°w 20°w 15°w 40°w 35°w 30°w 25°w 20°w 15°w 10°w 83 °n 82 °n 81 °n 80 °n 83 °n 82 °n 81 °n 80 °n 0 50 100 km 1848 m sea and lake ice 300–1500 >1500 elevation (m a.s.l.) 0–300 fig. 19 fig. 3 elevation and place names of the study area in eastern north greenland. pr: prinsesse. https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 5 of 43 www.geusbul let in.org land within the metasediments of the palaeozoic franklinian basin. sediments as young as late cretaceous (and possibly palaeogene) occur within the fault zone. 3. the north-west to south-east-trending trolle land fault z one (tlfz) is more than 300 km long and 100 km wide and consists of a number of linear faults that cut through sediments as young as late cretaceous. according to piepjohn & von gosen (2003), the ‘wandel hav mobile belt’ (their term) is bordered by the hffz in the north and the tlfz in the south-west. paech & estrada (2018) pointed out that the age relations between these fault zones and the eurekan deformation stages are still unclear. 2.1 stratigraphic record throughout this section, we refer the reader to the geological map and stratigraphic column for the wandel sea basin presented in figs. 4 and 5, respectively. 2.1.1 palaeozoic strata the fluvial deposits of the lower carboniferous sortebakker formation (thickness >1 km) represent the k r o n p r i n s c h r i s t i a n l a n d ø hovgaard p e a r y l a n d 48 10°w 83 °n 82 °n 81 °n 80 °n 15°w20°w25°w30°w35°w 40°w 15°w20°w25°w30°w35°w nakkehoved amdrup land kilen holm land pr.t.ø pr.m.ø kr depotbugtfrigg fjord herluf trolle land paleocene–eocene sediments upper cretaceous sediments kap washington group jurassic–cretaceous sediments triassic sediments carboniferous–permian sediments cambro–silurian sediments precambrian rocks sea and lake quaternary ice fault dykes, mostly upper cretaceous sedimentary outlier kctz hffz tlfz tlfz eg fz 100 km500 fig. 4 geology of the carboniferous–palaeogene wandel sea basin and surrounding areas. the remnants of the carboniferous to palaeogene sediments occur between northern peary land and kronprins christian land and thus define the extent of the basin. kap washington group is of campanian to earliest palaeogene age (håkansson & pedersen 2015). håkansson & pedersen (2015) indicated the trolle land fault zone as one continuous zone from herluf trolle land to kronprins christian land. here, we indicate one possible correlation between these two areas. based on escher & pulvertaft 1995 with modifications after croxton et al. 1980; hovikoski et al. 2018; piasecki et al. 2018. egfz: east greenland fracture zone. hffz: harder fjord fault zone. kctz: kap canon thrust zone. kr: kap rigsdagen. pr.m.ø: prinsesse margrethe ø. pr.t.ø: prinsesse thyra ø. tlfz: trolle land fault zone. https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 6 of 43 www.geusbul let in.org ma era stage/ age eocene oligocene miocene pliocene pleistocene series/ epoch system/ period middle middle lower upper guadalupian lopingian lower lower middle mississippian upper mississippian lower penns. middle penns. upper penns. cisuralian lower mississippian langhian aquitanian chattian rupelian priabonian bartonian lutetian thanetian selandian danian maastrichtian campanian santonian coniacian aptian hauterivian valanginian berriasian tithonian bathonian bajocian aalenian pliensbachian sinemurian hettangian carnian ladinian anisian olenekian induan rhaetian roadian wordian capitanian wuchiapingian changhsingian kungurian artinskian asselian gzhelian kasimovian moscovian bashkirian serpukhovian visean tournaisian barremian toarcian ypresian burdigalian pa la eo zo ic m es o zo ic c en o zo ic n eo g en e pa la eo g en e c re ta c eo u s ju ra ss ic tr ia ss ic pe rm ia n c a rb o n if er o u s sakmarian upper kimmeridgian oxfordian albian cenomanian turonian serravallianserravallian tortonian messinian upper norian paleocene callovian 0 10 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 210 220 230 240 250 260 270 280 290 300 310 320 330 340 350 kw peary land k.c. land wandel sea basin is ra nd kap jungersen pb sortebakker mågensfjeld herlufsholm strand la de gå rd så en sølverbæk ? ? thyra ø nakkehoved channel fills marine mudstone marine carbonate marine sandstone sedimentary environments lithological symbols coal layers sandstone siltstone mudstone/shale organic rich mudstone/shale limestone fluvial, lacustrine extrusive igneous rocks volcanics, volcanoclastics evaporites conglomerate foldedal kim fjelde lichenryg dunken storekløft midnatsfjeld g al ad ri el f je ld legend bf dd kl fig. 5 wandel sea basin stratigraphy and formation names. the representation of the mixed volcanic, volcano-clastic and lacustrine deposits of the kap washington group has been simplified. some formation names are not shown; the upper permian kap kraka fm in peary land (håkansson & pedersen 2015) and the pre-late bajocian (?) gletscherport fm in the kilen area (hovikoski et al. 2018). compiled after bjerager et al. 2019; gautier et al. 2011; håkansson & pedersen 2015; hopper et al. 2014; hovikoski et al. 2018; ineson et al. 2020; piasecki et al. 2018; stemmerik et al. 1998; svennevig et al. 2018; tegner et al. 2011. bf: birkelund fjeld. dd: dromledome. k.c. land: kronprins christian land. kl: kuglelejet. kw: kap washington group. pb: parish bjerg. penns: pennsylvanian. oldest unit of the wandel sea basin. it crops out on holm land, where it rests directly on the crystalline basement affected by the caledonian orogeny (stemmerik et al. 1998). the upper carboniferous kap jungersen and foldedal formations show a marked variation from laterally widespread mixed siliciclastic-limestone cycles to dolomitised limestone where they crop out on the eastern kronprins christian land. the foldedal formation is also exposed on peary land. above a lower permian hiatus, the middle to upper permian limestones and shales of the kim fjelde and the midnatfjeld formations are exposed on herluf trolle land and on eastern kronprins christian land. the upper permian, fluvial to lacustrine deposits with syndepositional volcanics of the kap kraka formation (>1.5 km; not shown) are located in the central part of hyde fjord, northern peary land, where the deposits are strongly compressed and bounded by the faults of the hffz (håkansson & pedersen 2015). unnamed permian carbonates and deep-marine shales (>2 km) are exposed on prinsesse ingeborg halvø, where they appear as thrust sheets bounded by two of the main faults of the tlfz (håkansson & pedersen 2015). 2.1.2 mesozoic strata the wandel sea basin contains a nearly complete triassic succession, with an estimated combined thickness https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 7 of 43 www.geusbul let in.org of 700 m (bjerager et al. 2019). the basal part of the lower triassic succession is fluvial, and the remaining triassic succession is marine, representing a sand-dominated basin marginal succession in peary land to the west and a marine basinal mudstone succession in kronprins christian land to the east. the succession includes the lower triassic parish bjerg formation that crops out in northern herluf trolle land where it rests unconformably on upper permian sediments. the triassic sediments on herluf trolle land are bounded by two of the main faults of the tlfz (håkansson & pedersen 2015). alsen et al. (2017) identified mainly the middle triassic isrand formation in the kilen area on kronprins christian land. this formation consists of laminated mudstones that accumulated in a slope and basin-floor setting in the eastern and deeper part of the wandel sea basin. these deposits accumulated at the same time as shallow-marine mudstones of the dunken formation on herluf trolle land (alsen et al. 2017). a sedimentary outlier in amdrup land is now dated as early triassic (alsen et al. 2017). a minor outcrop of middle jurassic (bajocian) muddy sandstones occurs within a few square kilometres on central herluf trolle land within a dense fault complex in the centre of a broad domal structure (håkansson & pedersen 2015). upper jurassic – lower cretaceous mainly marine sandstones and sandy shales of the ladegårdsåen formation (250 m) crop out on herluf trolle land. the upper ladegårdsåen formation is a sandstone-dominated unit deposited in a fluvial through estuarine to fully marine environment, and it rests unconformably on the marine sandstones of the lower ladegårdsåen formation (ineson et al. 2020). the lower part of the formation rests unconformably on palaeozoic rocks within several isolated, typically fault-bounded occurrences across most of the tlfz (håkansson & pedersen 2015). piasecki et al. (2018) identified a c. 75 m high sedimentary section exposed at kap rigsdagen as part of ladegårdsåen formation, overlain by 5–10 m of the palaeogene thyra ø formation (see section 2.1.3). unnamed lower cretaceous (albian–aptian), mainly marine mudstones (c. 100 m) bounded by faults, crop out within a few square kilometres on central herluf trolle land (håkansson & pedersen 2015). turonian to coniacian, fluvial to marginally marine sand-shale series of the herlufsholm strand formation (500 m) crops out on northern herluf trolle land. the sediments are bounded by the northernmost fault of the tlfz, and they rest unconformably on palaeozoic sediments (håkansson & pedersen 2015). the poorly dated upper cretaceous marine sandstone of the nakkehoved formation (>600 m) crops out on the north coast of kronprins christian land, near the northern margin of the tlfz (håkansson & pedersen 2015). the sandstone is gently folded, but intensely cut by a swarm of mono-mineralic quartz veins in the modestly deformed sandstone. the vein system is essentially confined to a small semi-nunatak in the ice cap (håkansson et al. 1994). hovikoski et al. (2018) argued that kilen contains the thickest and stratigraphically most complete jurassic and cretaceous sediment succession in north greenland and presented a new stratigraphic scheme. the jurassic succession (>500 m) is divided into four formations consisting of lagoon and shallow-marine sediments: the poorly dated gletscherport formation, the mågensfjeld and birkelund fjeld formations of bajocian–bathonian and kimmeridgian age, respectively, and the volgian kuglelejet formation. the lower cretaceous interval (>1500 m) is divided into three formations consisting of deep shelf to fluvial and shoreface sediments: the late ryazanian to hauterivian dromledome formation, the unfossiliferous lichenryg formation and the late aptian to middle cenomanian galadriel fjeld formation. the lichenryg formation and the contemporaneous upper ladegårdsåen formation in peary land show close stratigraphic and palaeo-environmental parallels: both overlie an hauterivian – early barremian, erosional unconformity and show clear transgressive trends from fluvial through estuarine to fully marine (ineson et al. 2020). ineson et al. (2020) interpreted the succession to record hauterivian – early barremian regional uplift and erosion followed by fluvial sedimentation and subsequent transgression in the late barremian – early aptian. it was suggested that the uplift heralded the onset of the high arctic large igneous province (halip) at about 135 ma and marked the initiation of the amerasia basin (døssing et al. 2013). the upper cretaceous marine succession (650 m) is assigned to the late cenomanian to santonian sølverbæk formation dominated by mudstones and sandstone-mudstones. intense compressive deformation affected the mesozoic deposits in kilen (håkansson & pedersen 2001, 2015; svennevig et al. 2016, 2017). unnamed, deep-marine, siliciclastic sediments (>400 m) of santonian age occur in the frigg fjord area, bounded by the hffz, and their near-vertical orientation indicates substantial shortening (håkansson & pedersen 2015). the sediments are intruded by doleritic dikes and sills of similar composition to the basalts of the kap washington group, and these intrusions are affected by deformation along the hffz (paech & estrada 2018). further east along the hffz, about 500 m of upper cretaceous, mainly fluvial sandstone occur in https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 8 of 43 www.geusbul let in.org the depotbugt area (håkansson & pedersen 2015). their similarity to the herlufsholm strand formation has been used to date these sediments. volcanics and volcanoclastics interbedded with fluvial and lacustrine deposits of the campanian to lowermost palaeogene kap washington group (>5 km) occur on the north coast of peary land (håkansson & pedersen 2015). most of the volcanics were emplaced at 71–68 ma but activity continued into the paleocene (tegner et al. 2011). the deposits, which overlie palaeozoic strata, are bounded on all sides by the thrusts of the kctz, interpreted as reactivated normal faults (håkansson & pedersen 2015). 2.1.3 palaeogene strata the fine-grained sandstones, siltstones and coal of the thyra ø formation that crop out on prinsesse thyra ø, prinsesse dagmar ø, prinsesse margrethe ø and prinsesse ingeborg halvø are dated as late paleocene to possibly earliest eocene age, with an estimated composite thickness of 50 m (lyck & stemmerik 2000). håkansson et al. (1991) interpreted the depositional environment as dominantly fluviatile; however, lyck & stemmerik (2000) documented a marine influenced environment. in contrast to the folded and thrusted upper cretaceous and older deposits within the hffz and tlfz (e.g. in kilen), the deposits of the thyra ø formation are often reported as undeformed (håkansson & pedersen 1982, 2015; boyd et al. 1994; lyck & stemmerik 2000), but some deformational structures are also reported (pedersen & håkansson 2001; piepjohn & von gosen 2001; guarnieri 2015). the base of the formation is not known (lyck & stemmerik 2000) and on prinsesse ingeborg halvø, the palaeogene deposits are in contact with late paleozoic sediments along a wide fault zone (see section 2.2). minor outliers of palaeogene carbonaceous, terrestrial shales occur in the depotbugt area near the upper cretaceous outcrops (håkansson & pedersen 2015; paech & estrada 2018). croxton et al. (1980) dated a sparse pollen flora from these deposits as eocene in age, and paech & estrada (2018) noted that they were flat lying. piasecki et al. (2018) identified the upper 5–10 m of the sedimentary section exposed at kap rigsdagen to be of early to middle eocene age and to be part of the thyra ø formation, above sediments of ladegårdsåen formation separated by a 65 myr hiatus. abundant, reworked cretaceous dinoflagellate cysts are common in the thyra ø formation (lyck & stemmerik 2000; piasecki et al. 2018). extensive reworking shows that cretaceous units were exposed and actively eroded during deposition of the thyra ø formation (piasecki et al. 2018). 2.2 structural elements of the wandel sea basin and their relation to the eurekan orogeny in this section, we discuss the dominant structural elements of the wandel sea basin and their relation to the eurekan orogeny. we use the term eurekan orogeny to mean deformation caused by the palaeogene convergence between greenland, the canadian arctic and svalbard (figs. 1, 2). we use the development of the seafloor spreading in the north-east atlantic to define the following two stages of the eurekan orogeny: 1. sea-floor spreading in the labrador sea and baffin bay started in the mid-paleocene (magnetochron c27; c. 62 ma). the movement of greenland relative to the canadian arctic and svalbard resulted in the first stage of the eurekan orogeny (phase 1 of gion et al. [2017] and phase 2 of oakey & chalmers [2012]). 2. greenland and eurasia moved together until the latest paleocene (before c24, c. 55 ma; gaina et al. 2017), when the sea-floor spreading started between greenland and europe. subsequently, greenland moved north relative to the arctic islands, which resulted in the second stage of the eurekan orogeny (phase 2 of gion et al. [2017] and phase 3 of oakey & chalmers [2012]). sea-floor spreading between greenland and north america ceased at the end of the eocene (c13, c. 35 ma), when greenland became attached to north america and the eurekan orogeny ended. phases 1 and 4 of oakey & chalmers (2012) consist only of continental extensional movements between greenland and north america prior to the mid-paleocene and of continued movement of the new greenland – north america plate relative to eurasia after the eocene. none of the movements during these phases are, here, attributed to the eurekan orogeny. initial movement during the paleocene led to the formation of foreland basins, for example, the central tertiary basin (ctb) on svalbard, and to initial transpression on ellesmere island, where there was also some compression on at least one step-over between strike-slip faults during the first eurekan stage, defined above. a subsequent change of direction of greenland relative to north america during the eocene resulted in compression on ellesmere island and transpression on svalbard during the second eurekan stage (steel et al. 1985; ricketts 1994; harrison et al. 1999; faleide et al. 2008; oakey & chalmers 2012; dallmann 2015; embry & beauchamp 2019). the svalbard margin evolved through strike-slip movements along the plate boundary, the de geer line, where greenland slid past svalbard (steel et al. 1985; https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 9 of 43 www.geusbul let in.org faleide et al. 2008). the eurekan orogeny came to an end when the movement ceased between greenland and north america at the end of the eocene (oakey & chalmers 2012; piepjohn et al. 2016). sea-floor spreading between greenland and svalbard began in the early miocene (jokat et al. 2016). håkansson & pedersen (2015) observed a 45° offset in the plate boundary between north greenland and svalbard relative to the main trend of rifting and spreading in the north-east atlantic. as a result, this segment of the plate boundary experienced episodes of combined transtension and transpression, in part controlled by the movement of a temporarily independent greenland plate. they concluded that the upper permian – mesozoic deposits of the wandel sea basin record the plate-boundary history along this offset in a series of disturbed, pull-apart basins. they assigned most of these basins to four tectonic episodes that were finally overprinted by the compressional kronprins christian land orogeny (fig. 2). according to these authors, the timing of this orogeny is bracketed by the age of the deformed strata (upper cretaceous) and the age of the undeformed cover successions (thyra ø formation), and thus, occurred during a comparatively brief period around the cretaceous–palaeogene boundary. the initial formation of the long, linear faults of the tlfz was related to deformational events in late palaeozoic, triassic–jurassic and late cretaceous times (von gosen & piepjohn 2003). piepjohn & von gosen (2001) observed dextral transpressive deformation, coeval with north–south compression at the hffz and kctz, which they interpreted as eurekan. von gosen & piepjohn (2003) argued that the eurekan transpressive deformation of the wandel sea basin was younger than the thyra ø formation and thus of post-paleocene age, probably eocene. they interpreted a dextral displacement of the thyra ø formation relative to upper palaeozoic strata within a wide fault zone on prinsesse ingeborg halvø to have occurred in the early eocene during the deformation of the hffz and kctz. in contrast, pedersen & håkansson (2001) explained the displacement of the thyra ø formation along the fault in terms of post-paleocene down-faulting after cessation of compression. håkansson & pedersen (2001, 2015), therefore, suggested that the compressive tectonics had come to an end before the deposition of the thyra ø formation. guarnieri (2015) identified compressive structures and reverse fault planes along the contact on prinsesse ingeborg halvø and interpreted them as a paleocene thrust fault, offset by strike-slip faults of probable eocene. to our knowledge, prinsesse ingeborg halvø is the only locality where the sediments of the thyra ø formation are observed in contact with older rocks, and thus, the timing of the end of compressional tectonics remains to be speculative. svennevig et al. (2016, 2017) interpreted the cretaceous succession exposed on kilen as a rift basin, dominated by post-coniacian extensional faults and overprinted during north–south compression that possibly occurred during the paleocene–eocene. svennevig et al. (2016) also presented a new structural model for the development of the tlfz based on observations from kilen. previous structural models interpreted rhombic-shaped fault patterns as evidence for strike-slip tectonics. however, these authors demonstrated the structures to be the result of post-coniacian, north-east to south-west extension defined by north-west to south-east-trending normal faults, followed later by north–south compression that folded the faults passively. consequently, the strike-slip faults had only minor status in the resulting basin inversion. svennevig et al. (2016) considered the age of the compressional event on kilen to be post-coniacian, since the youngest cretaceous outcrops are coniacian in age. hovikoski et al. (2018) dated these outcrops as santonian. they also thought that a late paleocene – early eocene age for the north–south compression on kilen was found to be compelling because (1) the ar-ar systematics of volcanics of the kap washington group (200 km northwest of kilen) were partially reset in the early eocene (tegner et al. 2011), (2) the west spitsbergen fold belt on the conjugate barents margin has been dated to be late paleocene-eocene in age (maher et al. 1995) and (3) plate reconstructions showed transpression with 18° convergence in the late paleocene (60 ma) that culminated in the early eocene (54–50 ma; svennevig et al. (2016) based on müller et al. 2016). 3 apatite fission-track analysis a variety of scenarios have been proposed in the literature to explain the complex tectonic development of the wandel sea basin. studies in the region describe multiple episodes of folding and faulting, although the timing remains uncertain. here, we present afta data and thermal history interpretations, with the aim of placing firmer constraints on the timing of key tectonic events. afta is a method for determining thermal histories of rocks at temperatures generally less than 130°c. thermal history constraints derived from afta can be converted into information on former depths of burial and amounts of exhumation. in the method, radiation damage features known as ‘fission tracks’ are analysed. fission tracks form in accessory apatite crystals separated from igneous rocks or in detrital apatite grains obtained from sandstones and other clastic rock types. the number of tracks per unit area of a polished grain surface depends on the uranium content, the time over which tracks have accumulated and the distribution of track lengths in the sample. if all tracks have the same length, then by measuring the https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 10 of 43 www.geusbul let in.org track density and the uranium content, a ‘fission-track age’ can be measured. in the absence of other factors, this age represents the time over which tracks accumulated. once formed, the damage is progressively repaired, and the tracks shorten at a rate dependent on temperature. the probability of intersecting a polished surface depends on the track length, and shortened tracks produce a lower contribution to the track density than longer tracks, resulting in a reduced fission-track age. as a result, a measured fission-track age does not denote the timing of a specific event and must be assessed together with the distribution of track lengths, referred to as ’confined tracks’. these tracks are totally enclosed within the grains and have been etched because they intersect tracks etched from the surface. the principles of the afta method are further described in japsen et al. (in press), green & duddy (2012) and green et al. (2013). 3.1 new afta data from northern and eastern greenland this research study presents new afta data from 52 outcrop samples of basement and sedimentary rocks in northern and eastern greenland (fig. 6; tables 1, 2 and the supplementary file s1). of these, 46 are from eastern north greenland (north of 79°n; east of 40°w) and 6 are from north-east greenland (between 74 and 79°n). sixteen further samples were processed but did not yield apatite. as far as we are aware, no apatite fission-track data have yet been reported from greenland north of 80.5°n. methods for sample preparation and analysis, basic afta data and thermal history interpretations derived from afta data for all samples were first reported in geotrack report gc1113 (green 2014) – a report prepared by geotrack international, australia, for the geological survey of denmark and greenland, as provided in supplementary file s2. in brief, rock samples are crushed to sand grade and heavy minerals are separated using standard density and magnetic techniques. apatite grains are mounted in epoxy on glass slides, polished and etched in dilute nitric acid. the etched grain mounts are then fixed to mica sheets and irradiated with thermal neutrons. after irradiation, the mica sheets are detached from the grain mount and etched in hydrofluoric acid. finally, the etched grain mounts and mica sheets are mounted next to each other on a microscope slide for track counting and length measurement. ø hovgaard 48 83 °n 82 °n 81 °n 80 °n 15°w 10°w20°w25°w30°w35°w40°w 15°w20°w25°w30°w35°w a gc1113 samples 100 km500 39 37 43 1544 16 4532 3 30 29 33 34 36 38 6 17 10 64 63 59 65 23 68 67 21 22 5 42 54 1 25 58 8 60 62 18 24 1948 61 7 9 55 3 5 24 1 eg fzk r o n p r i n s c h r i s t i a n l a n d p e a r y l a n d tlfz hffz kctz nakkehoved h e r l u f t r o l l e l a n d kr pr.t.ø kilen tlfz tlfz freuchen land land nansen p. koch fjord 45°w 35°w 25°w 15°w 5°w 83 °n 80 °n 77 °n 25°w 4950 51 53 46 52 48 15 0 50 100 km 6 see fig. 6a b fig. 6 location of samples that yielded apatite and division of the study area into six geographical areas with similar thermal histories used to define the timing of key episodes in fig. 10. a: areas 1–5. b: area 6. geology legend in fig. 4. egfz: east greenland fracture zone. hffz: harder fjord fault zone. kctz: kap canon thrust zone. kr: kap rigsdagen. pr.t.ø: prinsesse thyra ø. tlfz: trolle land fault zone. https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 11 of 43 www.geusbul let in.org table 1 sample details. a digital version of this table is available in online supplementary file s1 sample number gc1113stratigraphic age chronstratigraphic age (ma) rock type latitude longitude elevation (m a.s.l.) location 1 coniacian 89–85 conglomerate –14.0866 81.3477 250 kcl, kilen 3 cambrian 542–488 quartzite –40.1618 83.2950 465 lockwood ø 5 cretaceous, l 100–65 sandstone –21.6885 82.7712 110 htl, foldedal 6 cambrian, e 542–513 sandstone –31.4713 82.1574 114 heilprin land, glaciologelv 7 cambrian, e 542–513 sandstone –20.6893 80.2133 160 pr. cm alper 8 carboniferous, e 360–318 sandstone –15.1754 80.7875 100 al, sommerterrassen 9 palaeogene, e 66–61 sand/coal –19.2820 81.9368 12 pr. thyra ø 10 jurassic, l – cretaceous, e 161–100 sandstone –21.2511 82.0323 23 kap rigsdagen 15 cambrian 542–488 psammite –34.8881 83.3972 1341 jvjl, polkorridoren 16 cambrian 542–488 marble? –35.6985 83.5279 320 jvjl, sands fjord 17 neoproterozoic 1600–1000 sandstone –26.2725 81.9563 275 iccl, neergaard dal 18 carboniferous, l 318–299 sandstone –15.7926 80.4417 20 hl, eskimonæs 19 carboniferous, e 360–318 sandstone –17.4994 80.2161 50 hl, sortebakker 21 oxfordian–barremian 161–130 sandstone –21.4981 82.5801 157 htl, kim fjelde 22 triassic, e 251–247 sandstone –21.3657 82.5852 200 htl, kim fjelde 23 triassic, e 251–247 sandstone –21.2864 82.6155 190 htl, falkefjeld 24 carboniferous, l 318–299 sandstone –16.2082 80.3018 37 hl, hanseeraq fjord 25 albian 112–100 sandstone –12.6921 81.1447 25 kpc, kilen 29 cambrian, ellesm. deform. 542–488 psammite –30.5000 83.3000 400 jvjl 30 cambrian, ellesm. deform. 542–488 –34.0000 83.5667 340 jvjl 32 cretaceous, l /palaeogene 80–50 rhyolite –39.6667 83.4333 410 jvjl 33 cambrian, ellesm. deform. 542–488 arkose –28.5000 83.2667 825 jvjl 34 cambrian, ellesm. deform. 542–488 psammite –27.0833 83.3000 460 jvjl 36 cambrian, ellesm. deform. 542–488 sandstone –26.1667 83.2833 20 jvjl 37 cretaceous, l /palaeogene 80–50 gabbro –32.7000 83.1833 950 jvjl 38 silurian 444–416 greywacke –34.2520 82.2700 775 pl 39 silurian 444–416 sandstone, partly –37.6167 83.9333 600 pl 42 palaeoproterozoic 2500–1600 sandstone –32.7500 81.7009 720 pl, vildtland 43 cambrian, ellesm. deform. 542–488 conglomerate –37.1000 83.2333 390 jvjl 44 cambrian, ellesm. deform. 542–488 psammite –36.0000 83.4333 1180 jvjl 45 cambrian, ellesm. deform. 542–488 schist –37.6333 83.5500 80 jvjl 46 precambrian, caled. deform. 2500–1600 metagabbro –19.5000 77.8000 30 gamma ø, nordmarken 48 precambrian, caled. deform. 2500–1600 gabbro –18.4012 79.7217 250 hovgaard ø 49 precambrian, caled. deform. 2500–1600 mica schist –25.1697 74.4372 1450 bl 50 precambrian, caled. deform. 2500–1600 mica schist –25.7823 74.4418 1305 bl 51 precambrian, caled. deform. 2500–1600 gneiss –23.6667 78.1000 900 kfviiil, bildsøe nunatakker 52 precambrian, caled. deform. 2500–1600 banded sequence –23.3167 78.5333 660 kfviiil, grønne nuuatak 53 precambrian, caled. deform. 2500–1600 gneiss –18.5500 78.5667 10 franske øer, kap bergendahl 54 palaeoproterozoic 2500–1600 siltstone –28.3333 81.4333 251 iccl 55 cretaceous, l 100–65 sandstone –13.1110 81.6990 276 kcl,nakkehoved 58 cretaceous, e 145–100 sandstone –13.9481 81.2638 264 kcl, dromledome 59 cretaceous, l 100–65 sandstone –21.6794 82.7600 130 htl 60 mesoproterozoic 1600–1000 metasediments –15.2836 80.8965 320 al 61 archaean >2500 gneiss –17.1175 80.4068 952 hl 62 mesoproterozoic 1600–1000 sandstone –17.7671 80.7579 997 al, tobias gletscher 63 silurian 444–416 sandstone –22.8773 82.4813 122 htl, vitskøl elv 64 oxfordian–barremian 161–130 sandstone –22.8718 82.4682 277 htl, vitskøl elv 65 triassic, m 245–237 sandstone –21.2546 82.6158 250 htl, hjulsporsdalen 67 triassic, m 245–237 sandstone –21.0833 82.6703 229 htl, dunken 68 oxfordian–barremian 161–130 sandstone –21.1756 82.6753 642 htl, dunken al: amdrup land; bl: bartholin land; hl: holm land; htl: herluf trolle land; iccl: ic christensen land; jvjl: jv jensen land; kcl: kronprins christian land; kfviiil: kong frederik viii land; pl: peary land; pr cm alper: prinsesse caroline matilde alper; pr thyra ø: prinsesse thyra ø; e: early; m: middle; l: late; caled.: caledonian; ellesm.: ellesmerian. https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 12 of 43 www.geusbul let in.org table 2 afta data. a digital version of this table is available in online supplementary file s1 sample number gc1113stratigraphic age (ma) a ρd b (106 tracks/ cm2) ρs b (106 tracks/ cm2) ρi b (106 tracks/ cm2) (pχ2)c (%) (no of grains) fission track age d (ma) mean track length e (μm) std. dev. f (μm) 1 89–85 1.209 (1998) 0.343 (138) 2.698 (1085) 38 (20) 29.2 ± 2.8 14.1 ± 0.8 (7) 2.05 3 542–488 1.215 (1998) 0.557 (143) 3.607 (926) <1 (20) 38.8 ± 4.9† 14.2 ± 0.4 (12) 1.51 5 100–65 1.222 (1998) 0.044 (25) 0.258 (148) 37 (20) 39.1 ± 8.5 14.7 ± 0.3 (8) 0.72 6 542–513 1.228 (1998) 1.807 (522) 2.552 (737) <1 (16) 165.9 ± 17.5† 12.7 ± 0.3 (27) 1.58 7 542–513 1.235 (1998) 1.120 (642) 1.950 (1118) 3 (20) 127.1 ± 9.6† 12.8 ± 0.2 (55) 1.64 8 360–318 1.241 (1998) 1.578 (568) 1.842 (663) 47 (20) 199.1 ± 12.6 12.8 ± 0.2 (72) 1.50 9 66–61 1.248 (1998) 0.969 (344) 1.375 (488) 9 (20) 165.6 ± 16.5† 12.8 ± 0.4 (44) 2.35 10 161–100 1.254 (1998) 0.542 (195) 1.036 (373) 3 (19) 127.0 ± 15.8† 11.2 ± 0.5 (15) 1.80 15 542–488 1.260 (1998) 0.334 (176) 2.786 (1469) 6 (20) 28.7 ± 2.4 14.0 ± 0.4 (24) 1.88 16 542–488 1.267 (1998) 0.037 (27) 0.134 (97) 99 (19) 66.7 ± 14.6 14.2 ± 0.5 (2) 0.64 17 1600–1000 1.273 (1998) 1.306 (637) 1.837 (896) 4 (20) 173.1 ± 14.3† 13.0 ± 0.2 (54) 1.37 18 318–299 1.280 (1998) 1.499 (599) 2.275 (909) 4 (18) 153.8 ± 11.8† 11.3 ± 0.3 (21) 1.28 19 360–318 1.292 (1998) 1.130 (833) 1.566 (1154) 68 (20) 175.0 ± 9.2 12.3 ± 0.2 (100) 1.71 21 161–130 1.292 (1998) 0.712 (288) 2.088 (845) <1 (20) 83.4 ± 9.3† 11.9 ± 0.6 (17) 2.28 22 251–247 1.299 (1998) 1.140 (594) 2.140 (1115) <1 (20) 127.0 ± 11.9† 12.1 ± 0.2 (89) 2.06 23 251–247 1.305 (1998) 0.839 (526) 2.699 (1692) <1 (20) 76.3 ± 7.0† 13.3 ± 0.3 (40) 1.91 24 318–299 1.312 (1998) 0.923 (790) 1.249 (1069) 3 (20) 183.8 ± 13.2† 12.8 ± 0.2 (100) 1.60 25 112–100 1.318 (1998) 0.477 (69) 2.701 (391) 53 (12) 44.1 ± 5.9 13.8 ± 0.5 (5) 1.19 29 542–488 1.125 (1885) 0.480 (280) 3.051 (1780) 2 (16) 33.3 ± 3.2† 13.4 ± 0.3 (24) 1.68 30 542–488 1.132 (1885) 0.052 (19) 0.142 (52) 88 (13) 78.2 ± 21.1 32 80–50 1.146 (1885) 0.694 (155) 2.551 (570) <1 (11) 65.3 ± 18.6† 12.6 ± 0.6 (6) 1.53 33 542–488 1.153 (1885) 0.953 (413) 4.022 (1744) <1 (20) 49.4 ± 5.0† 14.3 ± 0.2 (52) 1.39 34 542–488 1.160 (1885) 0.296 (36) 3.598 (437) <1 (9) 24.9 ± 6.6† 13.9 ± 1.6 (3) 2.72 36 542–488 1.167 (1885) 1.268 (79) 2.372 (148) <1 (9) 115.9 ± 28.0† 13.9 ± 0.1 (2) 0.16 37 80–50 1.181 (1885) 3.009 (1367) 3.306 (1502) 2 (20) 201.9 ± 11.5† 13.5 ± 0.2 (106) 1.63 38 444–416 1.188 (1885) 2.790 (1415) 2.622 (1330) <1 (20) 237.3 ± 14.7† 13.1 ± 0.2 (100) 1.57 39 444–416 1.195 (1885) 1.282 (426) 1.664 (553) <1 (18) 164.0 ± 18.5† 13.1 ± 0.4 (18) 1.67 42 2500–1600 1.202 (1885) 3.017 (169) 3.357 (188) <1 (4) 174.5 ± 56.5† 14.1 ± 0.9 (5) 1.92 43 542–488 1.209 (1885) 0.818 (478) 4.463 (2609) <1 (20) 45.3 ± 6.0† 13.7 ± 0.2 (53) 1.31 44 542–488 1.216 (1885) 0.093 (8) 0.865 (74) 16 (3) 24.9 ± 9.3 14.2 ± 0.0 (1) 45 542–488 1.223 (1885) 0.877 (302) 1.984 (683) <1 (8) 90.0 ± 32.2† 12.0 ± 0.4 (10) 1.29 46 2500–1600 1.229 (1885) 0.913 (681) 1.163 (867) <1 (20) 163.5 ± 14.5† 13.4 ± 0.1 (100) 1.16 48 2500–1600 1.236 (1885) 1.953 (1563) 1.803 (1443) <1 (20) 251.7 ± 14.8† 13.4 ± 0.1 (101) 1.18 49 2500–1600 1.243 (1885) 0.986 (599) 1.625 (987) <1 (20) 136.1 ± 11.7† 13.2 ± 0.1 (77) 1.28 50 2500–1600 1.250 (1885) 1.982 (1545) 2.450 (1910) 2 (20) 184.7 ± 10.1† 13.4 ± 0.1 (110) 1.50 51 2500–1600 1.257 (1885) 1.479 (1264) 3.773 (3224) <1 (20) 93.9 ± 6.7† 13.4 ± 0.1 (101) 1.15 52 2500–1600 1.264 (1885) 0.759 (603) 1.375 (1092) 1 (20) 129.7 ± 9.8† 12.8 ± 0.2 (61) 1.35 53 2500–1600 1.178 (1924) 0.042 (38) 0.048 (43) 100 (18) 201.4 ± 45.2 13.0 ± 0.5 (15) 1.88 54 2500–1600 1.182 (1924) 1.986 (801) 1.539 (621) 3 (20) 293.7 ± 23.2† 13.1 ± 0.2 (106) 1.64 55 100–65 1.187 (1924) 0.343 (124) 1.838 (665) <1 (19) 48.6 ± 7.5† 11.7 ± 0.9 (10) 3.00 58 145–100 1.191 (1924) 0.552 (323) 2.747 (1606) 15 (20) 47.9 ± 4.1† 13.8 ± 0.2 (113) 1.95 59 100–65 1.195 (1924) 0.169 (118) 0.860 (601) 92 (20) 45.9 ± 4.8 13.6 ± 0.2 (67) 1.69 60 1600–1000 1.199 (1924) 0.478 (430) 0.424 (382) 32 (20) 259.8 ± 19.8 12.8 ± 0.1 (115) 1.40 61 >2500 1.203 (1924) 1.456 (766) 1.363 (717) 38 (20) 247.6 ± 14.8 12.8 ± 0.2 (109) 1.81 62 1600–1000 1.207 (1924) 2.156 (932) 1.855 (802) <1 (20) 258.2 ± 21.8† 13.0 ± 0.1 (116) 1.48 63 444–416 1.211 (1924) 2.129 (939) 1.757 (775) 2 (20) 276.8 ± 19.6† 12.0 ± 0.1 (108) 1.50 64 161–130 1.215 (1924) 2.559 (884) 2.255 (779) <1 (20) 299.1 ± 38.5† 11.8 ± 0.2 (105) 1.97 65 245–237 1.219 (1924) 0.941 (266) 2.297 (649) <1 (15) 80.6 ± 19.9† 13.2 ± 0.5 (22) 2.14 67 245–237 1.223 (1924) 0.979 (364) 4.609 (1714) 13 (20) 50.8 ± 3.3 14.0 ± 0.2 (105) 1.79 68 161–130 1.227 (1924) 0.682 (216) 3.545 (1122) 78 (20) 46.2 ± 3.7 13.5 ± 0.2 (119) 1.76 a all numerical values for stratigraphic ages assigned following gradstein et al. (2012). b ρs = spontaneous track density; ρi = induced track density; ρd = glass-dosimeter track density. numbers in parentheses show the number of tracks counted in determining all track densities. c probability that all singlegrain ages belong to a single population (galbraith 2005). d central age (galbraith 2005), used for samples containing a significant spread in single-grain ages (pχ2) < 5%), is denoted with †, otherwise the pooled age is quoted. all ages were calculated using the zeta calibration approach of hurford & green (1983), using zeta values for cn5 glass of 380.4 ± 5.7 (samples gc1113-1 to -52); 392.9 ± 7.4 (samples gc1113-53 to 68). all errors quoted at ± 1 σ.  all analytical details are as provided in supplementary file s1. e numbers in parentheses show the number of track lengths measured. f standard deviation of the track-length distribution. https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 13 of 43 www.geusbul let in.org the chlorine content exerts a critical influence on fission-track annealing kinetics in apatite. measuring chlorine is essential to obtain both improved accuracy and precision in the timing and magnitude of thermal events. chlorine compositions were determined by an electron microprobe (see supplementary file s2) for all individual apatite grains analysed (i.e. all grains in which fission-track ages were determined or lengths were measured). measured fission-track ages vary between 24.9 ± 6.6 ma and 299.1 ± 38.5 ma, with most samples (c. 80%) yielding ages less than 200 ma (table 2). apatite fission-track ages less than 300 ma in samples of precambrian basement or early palaeozoic metasediments show that the samples must have been much hotter than their present-day temperature at some time during the last 300 ma. youngest apatite fission-track ages (<50 ma) are generally found in coastal locations, while inland samples give consistently older ages (up to c. 300 ma; fig. 7). in addition to the thermal history, measured fission-track ages are sensitive to depositional ages, as well as external factors, such as the composition of apatite within each sample. thus, regional variation is better described by looking at palaeotemperatures within individual palaeo thermal episodes, as discussed in section 4.5. mean track lengths are generally between 12 and 14 mm, with a small number of samples giving higher values (towards c. 15 µm) and lower values (towards c. 11 µm; fig. 8a; table 2). initial assessment of these data suggests that they are dominated by the late palaeozoic and later history, and retain little trace of earlier events. the afta data define a broadly consistent trend of mean track lengths versus apatite fission-track ages (fig.  8a), where longest mean track lengths correspond to the youngest ages and the mean track length decreases as ages increase. while this trend resembles the classic ‘boomerang’ trend (green 1986), typical of a suite of samples affected by a single dominant palaeothermal episode, the data show much more dispersion, suggesting a more complex history involving a series of discrete cooling episodes. despite this complexity, the overall trend in fig. 8a suggests that most samples across the region have undergone a similar style of thermal history involving a series of cooling episodes (fig.  8a), with the magnitude of individual events varying across the region. extraction of thermal history information from the afta data and subsequent synthesis has, therefore, been carried out on this basis. as none of the precambrian samples give fission-track ages in excess of 300 ma, data from these samples were interpreted by reference to the base of the phanerozoic at 542 ma. afta data from the wandel sea basin show a similar range to those from north-east greenland (japsen et al. in press), in an area immediately south and overlapping with the study region (fig. 8b). however, the wandel sea basin samples plot consistently towards longer mean track lengths, particularly at younger ages. fig. 7 apatite fission-track ages in outcrop samples from eastern north greenland (data in table 2). youngest ages <60 ma are focussed along the north coast, while older ages are found in inland samples. egfz: east greenland fracture zone. hffz: harder fjord fault zone. kctz: kap canon thrust zone. kr: kap rigsdagen. pr.t.ø: prinsesse thyra ø. tlfz: trolle land fault zone. geology legend in fig. 4. ø hovgaard 48 83 °n 82 °n 81 °n 80 °n 15°w 10°w20°w25°w30°w35°w40°w 15°w20°w25°w30°w35°w 100 km500 eg fzk r o n p r i n s c h r i s t i a n l a n d p e a r y l a n d tlfz tlfz hffz kctz nakkehoved pr.t.ø kr kilen herluf trolle land fission-track age (ma) <25 25–45 45–60 60–100 100–200 >200 https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 14 of 43 www.geusbul let in.org this  probably reflects a reduced impact of neogene cooling in the samples analysed. this is consistent with lower summit heights in the study region, and therefore, less pronounced neogene uplift and incision. 3.2 thermal history interpretation based on afta data 3.2.1 principles of thermal history interpretation the techniques of extracting thermal history constraints from afta data are discussed in detail elsewhere (green & duddy 2012; green et al. 2013). in brief, fission tracks are formed continuously through time by spontaneous fission of 238u atoms and accumulate at a rate determined by radioactive decay laws. tracks form within a narrow range of lengths (between c.15 and 17 mm). as an apatite grain is heated, tracks are shortened at a rate that depends on the prevailing temperature. the kinetics of annealing are such that temperature dominates over time. in a sample that has been heated to a maximum palaeotemperature and then cooled, almost all of the tracks formed during the heating phase are shortened to the same length. if the sample remains at lower temperatures, tracks formed after cooling will subsequently undergo less shortening compared with those formed prior to the onset of cooling, resulting in two distinct populations of tracks. the mean length of the shorter population is determined by the maximum palaeotemperature, while the proportion of short vs. long tracks will be determined by the ratio of the time before and after the onset of cooling. if the sample is re-heated to a peak palaeotemperature that is lower than the previous peak, and then cools again, tracks formed prior to the maximum are unaffected but tracks formed after cooling from the earlier maximum will be shortened to a degree determined by the peak palaeotemperature in 9 10 11 12 13 14 15 16 0 50 100 150 200 250 300 350 m ea n tr ac k le ng th (µ m ) fission-track age (ma) a 9 10 11 12 13 14 15 16 0 50 100 150 200 250 300 350 m ea n tr ac k le ng th (µ m ) fission-track age (ma) this study ne greenland b fig. 8 relationship between the mean track length and fission-track age. a: outcrop samples from eastern north greenland (this study). error bars are shown as ± 2 sigma. b: results from this study compared with those from north-east greenland (japsen et al. in press). ages span a similar range in both regions. however, data from this study are characterised by higher mean track lengths than those from northeast greenland, which probably reflects the lower degree of miocene cooling in samples from this study. error bars are omitted for clarity. https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 15 of 43 www.geusbul let in.org the second heating event. this results in two generations of tracks with different lengths, which allows definition of two episodes of heating and cooling, if the two events are sufficiently separated in temperature and time. in some cases, three episodes might be resolved in a similar fashion, while definition of four episodes of cooling may be possible under rare circumstances. as afta data are dominated by the maximum palaeotemperature the data preserve no evidence of the history prior to the onset of cooling. for similar reasons, the history between successive cooling episodes cannot be resolved. therefore, it is not possible to constrain the entire thermal history from the beginning of track retention. instead, we define the maximum palaeotemperature and subsequent palaeothermal peaks within an overall framework of episodic heating and cooling (fig.  9), assuming heating and cooling rates of 1°c/ma and  10°c/ma, respectively. an order of magnitude change in heating rate is equivalent to around a 10°c difference in palaeotemperature required to produce a given degree of annealing. as explained in detail elsewhere, the combination of afta with geological evidence has provided abundant evidence that thermal histories involving episodic heating and cooling are more realistic than slow, monotonic cooling (green & duddy 2012; green et al. 2013, 2018). this conclusion is based on studies in which afta data provide evidence of cooling and exhumation, while times when underlying rocks were at the surface are identified from remnants of the former sedimentary cover. such situations provide evidence of repeated burial and exhumation, and are generally applicable in a wide variety of settings (green et al. 2013, 2017). thermal history solutions are represented by 95% confidence limits on the maximum (or peak) palaeotemperature and on the time at which cooling from that palaeotemperature began (fig. 9). the parameters are extracted from data in each sample by comparing measured afta parameters (fission-track ages and track lengths) with those predicted from candidate thermal histories, involving up to three episodes of heating and cooling. by varying the magnitude of maximum palaeotemperature and the onset of cooling in each episode, the range of temperature-time combinations giving predictions consistent with the measured data and within 95% confidence limits can be defined using the likelihood theory based on principles similar to those outlined by gallagher (1995). where three episodes are defined, the earliest episode is constrained primarily from the fission-track age data and represents the episode in which samples cool below c. 110°c and began to retain tracks, while the two more recent episodes are defined principally from the distribution of track lengths. 3.2.2 thermal history interpretation: results detailed thermal history interpretation of the afta data in each sample is provided in supplementary file s2. as explained therein, afta data in almost all samples show clear evidence of higher temperatures after deposition (or in phanerozoic times for samples of precambrian basement). thermal history solutions derived from the afta data in each sample, in terms of the maximum palaeotemperature and the timing of cooling in a number of discrete episodes, are summarised in table 3. fig. 9 illustration of thermal history solutions extracted from afta data. results are presented in terms of up to three palaeothermal episodes, that is, times when a rock sample was hotter than it is today. note that the palaeotemperature constraint t1 is a minimum estimate because this is the event in which the apatites begin to retain tracks. in most situations, three events are the most that can be defined from afta due to various factors, including the natural spread in track lengths of a single population of tracks and the rapid decrease in the rate of annealing across the range from c. 110°c to below 60°c. t1, t2 and t3: time intervals during which cooling from the peak palaeotemperature began. t1, t2 and t3: palaeotemperature intervals. c.i.: confidence interval. colours are used to illustrate attribution to regional episodes, although those used here are purely schematic. 100 80 60 40 20 0 120 250 200 50150 100 0 t2: 95% c.i. on maximum temperature t3 t3 t1 t1 t2: 95% c.i. on onset of cooling te m pe ra tu re (° c ) time (ma) deposition https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 16 of 43 www.geusbul let in.org ta bl e 3  th er m al h is to ry s ol ut io ns . a d ig ita l v er si on a va ila bl e in o nl in e su pp le m en ta ry f ile s 1 sa m pl e nu m be r g c1 11 3el ev at io n (m ) st ra tip ra ph ic a ge ch ro no st ra tigr ap hi c ag e (m a) pa la eo te m pe ra tu re (° c) o ns et o f co ol in g (m a) ev en t pa la eo – te m pe ra tu re (° c) o ns et o f co ol in g (m a) ev en t pa la eo – te m pe ra tu re (° c) o ns et o f co ol in g (m a) ev en t a re a 1                         32 41 0 cr et ac eo us , l /p al ae og en e 80 –5 0       11 0– 13 0 50 –2 0 en deo ce ne <1 10 20 –0 m io ce ne 37 95 0 cr et ac eo us , l /p al ae og en e 80 –5 0     85 –5 5 65 –1 0 en deo ce ne     3 46 5 ca m br ia n 54 2– 48 8     >1 05 45 –2 5 en deo ce ne     15 13 41 ca m br ia n 54 2– 48 8     >1 10 40 –2 5 en deo ce ne     16 32 0 ca m br ia n 54 2– 48 8     >1 00 12 0– 10 en deo ce ne     29 40 0 ca m br ia n, e lle sm . d ef or m . 54 2– 48 8 >1 10 75 –3 0 pa le oc en e 10 5– 12 0 45 –1 5 en deo ce ne 25 –9 5 25 –0 m io ce ne 30 34 0 ca m br ia n, e lle sm . d ef or m . 54 2– 48 8     95 –1 05 10 0– 0 en deo ce ne     33 82 5 ca m br ia n, e lle sm . d ef or m . 54 2– 48 8 >1 15 14 0– 60 pa le oc en e 11 0– 11 5 55 –3 5 en deo ce ne 20 –5 0 30 –0 m io ce ne 34 46 0 ca m br ia n, e lle sm . d ef or m . 54 2– 48 8 >1 05 >1 0 pa le oc en e >1 05 >1 0 en deo ce ne <1 20 20 –0 m io ce ne 36 20 ca m br ia n, e lle sm . d ef or m . 54 2– 48 8 ? ? pa le oc en e 10 0– 12 0 15 0– 20 en deo ce ne <1 00 11 0– 0 m io ce ne 43 39 0 ca m br ia n, e lle sm . d ef or m . 54 2– 48 8     >1 30 55 –4 0† en deo ce ne 40 –6 0 15 –0 m io ce ne 44 11 80 ca m br ia n, e lle sm . d ef or m . 54 2– 48 8     >1 00 70 –5 en deo ce ne     45 80 ca m br ia n, e lle sm . d ef or m . 54 2– 48 8       >1 15 50 –2 0 en deo ce ne       a re a 2                         5 11 0 cr et ac eo us , l 10 0– 65 >1 00 70 –2 0 pa le oc en e             59 13 0 cr et ac eo us , l 10 0– 65 >1 05 70 –4 0 pa le oc en e 55 –8 0 35 –1 0 en deo ce ne     10 23 ju ra ss ic , l – c re ta ce ou s, e 16 1– 10 0 >1 10 19 7– 13 6 l ju ra ss ic 81 –9 7 19 5– 2 en deo ce ne <7 8 46 –0 m io ce ne 21 15 7 o xf or di an –b ar re m ia n 16 1– 13 0 >1 00 >9 0 l ju ra ss ic 85 –1 00 55 –1 0 en deo ce ne     68 64 2 o xf or di an –b ar re m ia n 16 1– 13 0 >1 25 60 –4 0 pa le oc en e 90 –1 10 55 –3 5 en deo ce ne 40 –7 0 25 –5 m io ce ne 64 27 7 o xf or di an –b ar re m ia n 16 1– 13 0     75 –9 0 14 0– 25 en deo ce ne 15 –7 5 35 –0 m io ce ne 22 20 0 tr ia ss ic , e 25 1– 24 7 >1 15 20 0– 15 0 l ju ra ss ic 85 –9 5 75 –3 5 en deo ce ne 25 –5 0 20 –0 m io ce ne 23 19 0 tr ia ss ic , e 25 1– 24 7 10 0– 10 5 90 –6 0 pa le oc en e 65 –8 5 45 –1 5 en deo ce ne     67 22 9 tr ia ss ic , m 24 5– 23 7 >1 30 65 –4 5 pa le oc en e 85 –1 05 55 –3 5 en deo ce ne 30 –6 5 35 –5 m io ce ne 65 25 0 tr ia ss ic , m 24 5– 23 7 10 0– 10 5 12 5– 30 pa le oc en e 20 –9 5 60 –0 en deo ce ne     63 12 2 si lu ri an 44 4– 41 6 95 –1 05 36 0– 23 5 e pe rm ia n 70 –8 0 35 –0 en deo ce ne       a re a 3                         9 12 pa la eo ge ne , e 66 –6 1       20 –7 5 45 –0 en deo ce ne       9 pr ede p 12 >1 25 c 29 2– 18 5c l tr ia ss ic 91 –1 03 c 17 8– 80 c l ju ra ss ic 37 –8 0 93 –5 en deo ce ne 1 25 0 co ni ac ia n 89 –8 5     >1 30 45 –2 5 en deo ce ne     55 27 6 cr et ac eo us , l 10 0– 65 >1 25 10 1– 48 pa le oc en e 86 –1 06 56 –1 5 en deo ce ne 36 –8 1 20 –0 m io ce ne 25 46 al bi an 11 2– 10 0 >1 00 65 –1 0 pa le oc en e         58 26 4 cr et ac eo us , e 14 5– 10 0 >1 20 60 –4 5 pa le oc en e 76 –1 03 48 –2 6 en deo ce ne 25 –7 2 31 –8 m io ce ne co nt in ue d https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 17 of 43 www.geusbul let in.org sa m pl e nu m be r g c1 11 3el ev at io n (m ) st ra tip ra ph ic a ge ch ro no st ra tigr ap hi c ag e (m a) pa la eo te m pe ra tu re (° c) o ns et o f co ol in g (m a) ev en t pa la eo – te m pe ra tu re (° c) o ns et o f co ol in g (m a) ev en t pa la eo – te m pe ra tu re (° c) o ns et o f co ol in g (m a) ev en t a re a 4                         39 60 0 si lu ri an 44 4– 41 6 >1 10 >2 00 l tr ia ss ic 95 –1 10 20 0– –7 5 l ju ra ss ic 35 –7 0 45 –0 en deo ce ne 38 77 5 si lu ri an 44 4– 41 6 >1 05 29 5– 24 0 e pe rm ia n 65 –9 0 24 5– 30 en deo ce ne 20 –6 0 45 –0 m io ce ne 6 11 4 ca m br ia n, e 54 2– 51 3 >1 15 21 0– 16 0c l tr ia ss ic 80 –1 05 18 5– 75 l ju ra ss ic 45 –7 5 65 –0 en deo ce ne 17 27 5 n eo pr ot er oz oi c 16 00 –1 00 0 >1 15 23 0– 17 5 l tr ia ss ic 70 –9 5 18 0– 45 l ju ra ss ic 30 –6 5 40 –0 en deo ce ne 42 72 0 pa la eo pr ot er oz oi c 25 00 –1 60 0 >1 05 35 0– 16 5 e pe rm ia n         54 25 1 pa la eo pr ot er oz oi c 25 00 –1 60 0 >1 10 37 5– 29 0   70 –9 0 28 0– 70 l ju ra ss ic 25 –6 5 85 –0 en deo ce ne a re a 5                         18 20 ca rb on ife ro us , l 31 8– 29 9 >1 00 25 0– 17 0 l tr ia ss ic 80 –1 10 b 22 0– 0b l ju ra ss ic 80 –9 0 45 –0 en deo ce ne 19 50 ca rb on ife ro us , e 36 0– 31 8 >1 00 >1 75 l tr ia ss ic 80 –1 00 17 5– 75 l ju ra ss ic 50 –7 0 35 –0 en deo ce ne 24 25 ca rb on ife ro us , l 31 8– 29 9 >1 00 25 0– 20 0 l tr ia ss ic 75 –9 5 20 0– 65 l ju ra ss ic 30 –7 0 55 –0 en deo ce ne 8 10 0 ca rb on ife ro us , e 36 0– 31 8 >1 05 27 0– 20 0 l tr ia ss ic 75 –1 10 b 22 5– 65 b l ju ra ss ic 45 –7 5 50 –0 en deo ce ne 7 16 0 ca m br ia n, e 54 2– 51 3 >1 15 17 5– 13 0 ju ra ss ic 80 –1 05 15 5– 50 m id -c re t. 35 –7 5 60 –0 en deo ce ne 60 32 0 m es op ro te ro zo ic 16 00 –1 00 0 >1 00 35 5– 23 0 e pe rm ia n 75 –1 05 31 5– 14 5 l tr ia ss ic 55 –7 5 70 –1 5 en deo ce ne 62 99 7 m es op ro te ro zo ic 16 00 –1 00 0 >1 00 >2 50 e pe rm ia n 80 –1 05 30 0– 21 0 l tr ia ss ic 50 –6 5 60 –0 en deo ce ne 48 25 0 pr ec am br ia n, c al ed . d ef or m . 25 00 –1 60 0 >1 05 31 5– 25 5 e pe rm ia n     55 –7 0 70 –1 0 en deo ce ne 61 95 2 ar ch ae an >2 50 0 >1 00 >2 50 e pe rm ia n 85 –1 00 26 0– 16 5 l tr ia ss ic 55 –7 0 90 –2 0 en deo ce ne a re a 6                         46 30 pr ec am br ia n, c al ed . d ef or m . 25 00 –1 60 0 >1 10 23 0– 18 0 l tr ia ss ic 60 –9 0 18 5– 35 m id -c re t. 25 –6 0 30 –0 m io ce ne 49 14 50 pr ec am br ia n, c al ed . d ef or m . 25 00 –1 60 0     >1 00 18 0– 12 0 l ju ra ss ic 50 –6 5 35 –0 m io ce ne 50 13 05 pr ec am br ia n, c al ed . d ef or m . 25 00 –1 60 0 >1 05 22 5– 18 0   60 –7 0 90 –3 0 en deo ce ne     51 90 0 pr ec am br ia n, c al ed . d ef or m . 25 00 –1 60 0     >1 00 11 5– 90 m id -c re t. 50 –6 0 20 –0 m io ce ne 52 66 0 pr ec am br ia n, c al ed . d ef or m . 25 00 –1 60 0 >1 10 >1 65 l tr ia ss ic 10 0– 11 0 16 5– 10 0 l ju ra ss ic 55 –7 0 25 –0 m io ce ne 53 10 pr ec am br ia n, c al ed . d ef or m . 25 00 –1 60 0 >9 5 40 0– 55 l tr ia ss ic       35 –9 0 21 5– 0 m io ce ne a: o ut ly in g va lu e; b : u nc er ta in c on st ra in ts ; c : p re -d ep os tio na l c on st ra in ts ; e : e ar ly , m : m id dl e, l : l at e; c al ed .: ca le do ni an ; e lle sm .: el le sm er ia n; c re t.: c re ta ce ou s. ta bl e 3  th er m al h is to ry s ol ut io ns . a d ig ita l v er si on a va ila bl e in o nl in e su pp le m en ta ry f ile s 1 (c on tin ue d) https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 18 of 43 www.geusbul let in.org in most of the samples analysed, the afta data require at least three episodes of heating and cooling, while in a small number of samples only two episodes are required to explain the data. definition of two palaeothermal episodes represents a typical level of resolution that can be achieved from afta. generally, the earlier episode would be defined by the fission-track age data and possibly the shorter track lengths, while the more recent episode would be defined by the shortening of the main mode of the track length distribution. however, where the palaeotemperatures in each episode are sufficiently well separated, three episodes can be resolved. in samples gc1113-19 and -24, the earliest episode produces total annealing of all tracks, a later event leads to shortening of tracks formed after the initial cooling to lengths around 10–11 mm, and another later event then shortens tracks formed after the second episode to lengths around 13 mm. the ability to resolve three episodes often depends on the quality of the data, with 100 track length measurements normally required to provide a confident conclusion. in most samples from this study, this criterion was met. nevertheless, a random variation in measured parameters can lead to the failure of identifying three episodes in some samples, while nearby samples provide a clear resolution. thus, in some cases, the quoted thermal history solutions may represent the combined effects of unresolved multiple cooling episodes. this is also particularly important at low temperatures (<70°c), where differences in track length may be insufficient to allow multiple episodes to be resolved. synthesis of all data, as discussed in section 3.3, suggests that at least seven discrete cooling episodes are required to explain the results from all samples. this emphasises the fact that afta data in individual samples, even in the most favourable circumstances, allow resolution of only three discrete episodes. the events that dominate the data in a given sample will be a function of the magnitude of each episode, the quality of the afta data and the spread of compositions present. often, the highest cl levels may reveal an earlier episode, for which the evidence has been overprinted by later events in the more sensitive, low-cl apatites. 3.3 definition of major regional palaeothermal episodes figure 10 shows timing constraints (95% confidence limits) on the onset of discrete cooling episodes in individual samples defined by afta (fig. 9). thermal history constraints for all samples are listed in table 3. samples are grouped into six geographic regions based on the general uniformity of the data within each of the regions (areas 1–6; fig. 6). based on this uniformity of the data within each region and the evidence for regionally consistent thermal histories, we have sought to define the minimum number of regionally synchronous cooling episodes that can explain the results in all of these samples. integration of results from all samples requires a total of seven discrete cooling episodes (vertical bars in fig. 10; see also table 4). the onset of cooling in each episode ranges from early permian to middle to late miocene. these intervals represent the times at which cooling began in each episode, and we do not suggest that cooling was restricted to these intervals. note that in some cases, the estimated onset of cooling pre-dates the depositional age of the sample. this indicates cooling of the apatite in sediment provenance regions, prior to deposition in the host sample. pre-cenozoic episodes are expressed almost exclusively at inland locations (areas 4 and 5 together with locations to the south in area 6). three cenozoic episodes are focussed particularly in northern coastal locations (areas 1–3). two constraints shown in fig. 10 do not conform to the regional synthesis outlined above: (1) afta data in sample gc1113-43 (precambrian) define an initial phase of cooling from >110°c, which began between 55 and 40 ma. this falls between the two regional events that began at c. 60 ma and c. 35 ma. (2) the earliest cooling episode defined in sample gc1113-16 began between 210 and 160 ma, which is just outside the 225–210 ma timing derived from the majority of samples. these isolated examples most likely represent statistical outliers or may reflect the difficulty in resolving multiple cooling episodes within afta data in individual samples. further sampling in the region may help to resolve these issues. we have allocated the timing constraints for individual samples to specific regional events, such that adjacent samples that show cooling from similar palaeotemperatures with consistent timing constraints are interpreted as a common episode. we seek to minimise the total number of episodes required. in this way, samples with quite broad constraints can still be allocated to specific regional events. for example, samples gc1113-33 and -34 both cooled from palaeotemperatures >110°c, but while the timing constraint in sample –33 is quite well defined (140–60 ma), the timing in sample –34 is only defined as >10 ma. this is probably due to a combination of the low apatite yield (nine ages) and the young age (c. 25 ma), and as a result, only three track lengths were measured. given the proximity of both these samples, it seems reasonable to assume that both samples cooled in a common event, in this case the 60 ma cooling episode. this illustrates how data with quite broad constraints can still provide useful conclusions. the results of this study could be interpreted in many different ways, and attempting to explain the data in terms of the smallest number of regionally synchronous cooling episodes may mask a more complex regional variation in thermal history. however, the synthesis https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 19 of 43 www.geusbul let in.org fig. 10 timing constraints on the onset of cooling for discrete cooling episodes derived from the afta in each outcrop sample. thermal history solutions for each sample are listed in table 3. horizontal-shaded bars represent 95% confidence intervals on the time of cooling onset, as illustrated in fig. 9. samples are divided into six areas (areas 1–6; fig. 6). assuming these results represent regionally synchronous cooling episodes, results across the entire region require seven separate episodes, as shown by vertical shaded columns. constraints for the onset of each episode are listed at the bottom of the figure (see table 4). timing constraints for individual samples are assigned to a common regional event when adjacent samples cool from similar palaeotemperatures with consistent timings. in some cases, the estimated onset of cooling pre-dates the depositional age of the sample, indicating cooling of the apatites in sediment provenance regions prior to deposition in the host sample. 225–210295–290onset of cooling (ma): c. 35 15–8165–150 115–90 c. 60 gc1113-7 gc1113-1 gc1113-37 gc1113-32 gc1113-3 gc1113-5 gc1113-6 gc1113-9 gc1113-17 gc1113-23 gc1113-67 gc1113-65 gc1113-63 gc1113-24 gc1113-25 gc1113-58 gc1113-39 gc1113-38 gc1113-42 gc1113-54 gc1113-55 gc1113-21 gc1113-68 gc1113-64 gc1113-22 gc1113-10 gc1113-15 gc1113-16 gc1113-18 gc1113-60 gc1113-19 gc1113-29 gc1113-30 gc1113-33 gc1113-34 gc1113-36 gc1113-43 gc1113-44 gc1113-45 gc1113-59 area 1 area 2 area 3 area 4 area 5 area 6 gc1113-48 gc1113-46 gc1113-49 gc1113-50 gc1113-51 gc1113-52 gc1113-53 gc1113-61 gc1113-8 gc1113-62 050100150200250300350400 time (ma) stratigraphic age outlying value onset of cooling in indidual sample uncertain constraints onset of cooling across the region https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 20 of 43 www.geusbul let in.org presented here provides a consistent framework for interpreting the underlying nature of each episode. as discussed earlier, it is important to remember that afta data in individual samples only reveal two or at best three dominant episodes, and in at least some of the samples described here, the results probably represent the unresolved effects of a larger number of episodes. it is also important to realise that afta only constrains cooling from the palaeothermal peaks, and the thermal history in the interval between each cooling episode is not well constrained. a wide range of scenarios can, therefore, be invoked, involving cooling and reheating to a subsequent palaeothermal peak or prolonged stability between sequential cooling episodes. for all these reasons, definition of the thermal history in terms of a sequence of cooling episodes should be regarded as a general indication of the overall history, within the context of the cooling episodes defined from afta. 3.4 regional variation in palaeothermal episodes palaeotemperatures for six of the seven palaeothermal episodes identified are mapped in fig. 10, excluding samples from area 6 in fig. 6 (one episode is recognised only in a single sample from the region shown in fig. 11). here, we summarise the variation of palaeotemperatures in each episode and discuss the mechanisms responsible. early permian palaeotemperatures around 100°c or above, from which cooling began between 295 and 290 ma, are identified in samples scattered across most of the study area, but predominantly in the interior away from the major fault zones (fig. 11a). this is the earliest of the palaeothermal episodes recognised across the region, and as such, it is only recorded in areas where palaeotemperatures in subsequent episodes have remained sufficiently low as to not overprint it. locations recording this episode, therefore, represent relatively stable areas, which have undergone less post-permian heating and cooling than the surrounding regions. none of the samples in the region preserve evidence of remaining at temperatures below 100°c prior to the early permian times, emphasising the magnitude of post-palaeozoic palaeothermal effects across the entire region. late triassic palaeotemperatures, from which cooling began between 225 and 210 ma, occur predominantly in the centre of the study area and at south-eastern coastal locations (fig. 11b). maximum palaeotemperatures in this episode are in excess of 110°c for two samples in central locations and generally around 100°c or more on the south-eastern coast. as with the early permian episode, late triassic cooling is only recorded in areas where palaeotemperatures in subsequent episodes have remained sufficiently low as to not overprint evidence of this episode. late jurassic palaeotemperatures, from which cooling began between 165 and 150 ma, are more variable than those in earlier events (fig. 11c). two samples cooled from >110°c in this episode, which will have removed any evidence of earlier episodes. other samples cooled from palaeotemperatures generally between 80 and 100°c in this episode. the extent of this episode is similar in broad terms to the early permian and late triassic episodes but differs significantly from the three cenozoic episodes. only a single sample north of 79°n (fig. 6) shows definite evidence of mid-cretaceous cooling, which began between 115 and 90 ma (gc1113-17, not shown in fig. 11). however, this episode is also defined in samples to the south (area 6; fig. 10). evidence of cooling at this time is widespread in north-east greenland (japsen et al. in press). paleocene palaeotemperatures, from which cooling began at c. 60 ma (fig. 11d), are focussed almost exclusively within the hffz and the tlfz, as well as a single sample from kilen. end-eocene palaeotemperatures, from which cooling began at c. 35 ma, are recognised across almost the entire study area (fig. 11e). highest values (generally table 4 intervals defining the onset of episodes of cooling based on afta data in all samples and the possible origin of the cooling during each episode onset of cooling (ma) onset of cooling (chronostratigraphy) mechanism 295–290 early permian regional uplift and erosion 225–210 late triassic regional uplift and erosion 165–150 late jurassic regional uplift and erosion 115–90 mid-cretaceous regional uplift and erosion ~60 mid-paleocene uplift and erosion related to compression in fault zones ~35 end-eocene regional uplift and erosion, locally enhanced heating 15–5 middle to late miocene regional uplift and erosion figure 10 displays the timing of onset of discrete cooling episodes in individual samples. supplementary file s1 contains digital versions of table 1 (sample details and additional information), table 2 (afta data) and table 3 (thermal history solutions). supplementary file s2 contains geotrack report with analytical details, basic afta data in individual samples and a discussion of thermal history solutions derived from afta data in each sample (green 2014). https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 21 of 43 www.geusbul let in.org >110°c) are recognised not only in peary land, north of the hffz, but also at coastal locations south of the wandel sea basin; lower values around 60–70°c are identified across the interior of the region. the dominance of this episode, and to some extent also the paleocene episode, most likely masked the earlier episodes over much of the coastal region. the permian and triassic episodes are only recognised in regions where palaeotemperatures in palaeogene events were sufficiently low. on this basis, we suggest it is likely that these earlier episodes were also felt across the wandel sea basin, similar to events recognised in north-east greenland (japsen et al. in press). middle to late miocene palaeotemperatures, from which cooling began between 15 and 5 ma, are recognised mainly at coastal and near-coastal locations (fig. 11f), and generally at locations where palaeotemperatures at 35 ma were around 100°c or more. in samples that cooled from lower palaeotemperatures at 35 ma, it is likely that by 10 ma, these locations had cooled to temperatures too low to be resolved from afta data. 4 thermal history insights from maturity data a number of published studies provide further insights into the thermal history of the region. these are focussed mainly on thermal maturity studies, with some additional information provided by studies of diagenetic and mineralogical properties of rocks. here, we review these published data in light of the regional thermal history framework derived from afta in the previous section. the various datasets combined provide a broadly consistent picture of the thermo-tectonic development of the region. 4.1 maturity of lower palaeozoic strata rasmussen & smith (2001) sampled conodonts within the lower palaeozoic rocks across kronprins christian land (79–81°n) at the northernmost limit of the east greenland caledonides. they interpreted the conodont colour alteration index (cai) to provide a measure of maximum burial temperature and, in turn, the thickness of overburden. the data showed a gradual and continuous increase across kronprins christian land from cai 3 (110–200°c) in the west to cai 5+ (>300°c) in the east (armstrong et al. 1994). they concluded that the predominant process defining the regional variation was loading by caledonian thrust sheets. their modelling of the thickness of the removed overburden, based on a palaeogeothermal gradient of 28°c/km, suggested that it increased from 3.9 km in the west to a maximum of 12.5 km beneath a thrust sheet in the easternmost part of the area. rasmussen & håkansson (1996) reported a cai value of 6 for one sample of permian limestones from prinsesse ingeborg halvø (360-550°c). the sample may thus have reached these palaeotemperatures at any time since the permian. paech & estrada (2018) reported vr levels of 6.1-8.5% in rocks of probable ordovician age in the frigg fjord and kap washington regions (fig. 12), which indicate maximum palaeotemperatures >300°c. our afta data provide no control on pre-permian events, so no direct comparison with pre-permian palaeothermal episodes is possible. however, early permian cooling from >110°c (fig. 11a) would be consistent with palaeotemperatures in excess of 110°c in the early palaeozoic. 4.2 maturity of mesozoic–palaeogene strata published vr data for the wandel sea basin region are summarised in fig. 12. the data include (1) a comprehensive study of the thermal maturity of mesozoic and palaeogene sediments from the wandel sea basin, which overlaps with our study area (paech & estrada 2018); (2) a review of vr data (håkansson et al. 1994) and (3) a recently published study of vr data from kilen ( pedersen et al. 2018) that were not included in the review by paech & estrada (2018). maximum palaeotemperatures equivalent to the vr values in fig. 12 are calculated based on the algorithm of burnham & sweeney (1989). paech & estrada (2018) observed that mesozoic sequences attain high-maturity levels (vr mostly >2%; equivalent to >200°c) within prominent fault zones, including the kctz, hffz and tlfz (fig. 12). these higher maturity levels are typically associated with an increase in deformation intensity evidenced by the occurrence of a weak cleavage. in many cases, sediments as young as late cretaceous show evidence of profound tectonism, including steep dips and faulting. paech & estrada (2018) also noted that the high vr (up to 5.4%) levels for the kap washington group is associated with ductile deformation and magmatic activity. the maturity of jurassic–cretaceous sediments away from the major fault zones is relatively low, typically around 0.55% (c. 90°c), interpreted as reflecting heating due to regional burial. relatively flat-lying and undeformed palaeogene units on, for example, prinsesse thyra ø and in depotbugt (croxton et al. 1980; håkansson & pedersen 2015), contrast markedly with nearby, often intensely deformed upper cretaceous units that variously show steep dips, folding, faulting and, in some cases, thrusting. this contrast is supported by a major difference in maturity levels, with vr values around 0.55% in palaeogene units and 2% or more in cretaceous units within the major fault zones. paech & estrada (2018), therefore, concluded that the major phase of tectonism that resulted in such high maturity levels occurred around the cretaceous–cenozoic boundary. however, the https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 22 of 43 www.geusbul let in.org ø hovgaard 48 k r o n p r i n s c h r i s t i a n l a n d p e a r y l a n d tlfz tlfz hffz kctz holm land a palaeotemperature (°c) ø hovgaard 48 k r o n p r i n s c h r i s t i a n l a n d p e a r y l a n d tlfz hffz kctz tlfz holm land b palaeotemperature (°c) 10°w15°w20°w25°w30°w35°w40°w 83 °n 82 °n 81 °n 80 °n 10°w15°w20°w25°w30°w35°w40°w 83 °n 82 °n 81 °n 80 °n early permian (295–290 ma) late triassic (225–210 ma) fig. 11 palaeotemperatures derived from afta. a: early permian (295–290 ma). b: late triassic (225–210 ma). c: late jurassic (165–150 ma). d: paleocene (c. 60 ma). e: end-eocene (c. 35 ma). f: middle to late miocene (15–5 ma). palaeotemperature for the mid-cretaceous episode (115–90 ma) is not shown as only one sample falls within the map frame. the small arrow in the lower right corner of a, e and f indicates sample gc1113-48, which is located just south of the figure frame. geology legend in fig. 4. numbers next to afta samples denote sample numbers. hffz: harder fjord fault zone. kctz: kap canon thrust zone. kr: kap rigsdagen. pr.t.ø: prinsesse thyra ø. tlfz: trolle land fault zone. maturity of the palaeogene units is very similar to those in mesozoic units away from the fault zones, and palaeogene units do not occur within the zones of intense tectonism. it is, therefore, possible that the regional base level of maturity reflects heating after deposition of the thyra ø formation. extremely high maturity levels (>7%) were reported in the upper cretaceous nakkehoved formation exposed on the north coast of kronprins christian land (håkansson et al. 1994), corresponding to maximum palaeotemperatures well above 250°c (the limit at which palaeotemperatures can be calculated by vr). https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 23 of 43 www.geusbul let in.org ø hovgaard 48 k r o n p r i n s c h r i s t i a n l a n d p e a r y l a n d tlfz hffz kctz tlfz holm land herluf trolle land c palaeotemperature (°c) ø hovgaard 48 k r o n p r i n s c h r i s t i a n l a n d p e a r y l a n d tlfz hffz kctz nakkehoved kilen herluf trolle land tlfz d palaeotemperature (°c) 10°w15°w20°w25°w30°w35°w40°w 83 °n 82 °n 81 °n 80 °n 10°w15°w20°w25°w30°w35°w40°w 83 °n 82 °n 81 °n 80 °n paleocene (c. 60 ma) late jurassic (165–150 ma) quartz veins containing fluid inclusions with homogenisation temperatures of 197–236°c were formed during the thermal event, which was interpreted to have been short-lived and to have occurred at shallow burial depths (håkansson et al. 1994). samples of the nakkehoved formation showed diagenetic clay assemblages defining a low-grade greenschist metamorphic fabric, consistent with heating to palaeotemperatures of 200–300°c or above. palynology studies also indicate a high degree of thermal alteration, consistent with such palaeotemperatures. håkansson et al. (1994) found it unlikely that the thermal influence came from magmatic activity, in part due to the total absence of accessory minerals in the quartz veins. paech & estrada fig. 11 (continues) palaeotemperatures derived from afta. a: early permian (295– 290 ma). b: late triassic (225–210 ma). c: late jurassic (165–150 ma). d: paleocene (c. 60 ma). e: end-eocene (c. 35 ma). f: middle to late miocene (15–5 ma). palaeotemperature for the mid-cretaceous episode (115–90 ma) is not shown as only one sample falls within the map frame. the small arrow in the lower right corner of a, e and f indicates sample gc1113-48, which is located just south of the figure frame. geology legend in fig. 4. numbers next to afta samples denote sample numbers. hffz: harder fjord fault zone. kctz: kap canon thrust zone. kr: kap rigsdagen. pr.t.ø: prinsesse thyra ø. tlfz: trolle land fault zone. https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 24 of 43 www.geusbul let in.org ø hovgaard k r o n p r i n s c h r i s t i a n l a n d p e a r y l a n d tlfz hffz kctz nakkehoved kilen kr herluf trolle land pr.t.ø tlfz holm land e palaeotemperature (°c) ø hovgaard k r o n p r i n s c h r i s t i a n l a n d p e a r y l a n d tlfz hffz kctz nakkehoved kilen herluf trolle land f palaeotemperature (°c) 10°w15°w20°w25°w30°w35°w40°w 83 °n 82 °n 81 °n 80 °n 10°w15°w20°w25°w30°w35°w40°w 83 °n 82 °n 81 °n 80 °n end-eocene (c. 35 ma) middle to late miocene (15–5 ma) (2018) attributed the intense heat to the development of a plate boundary between north-eastern greenland and svalbard since the oligocene. however, why that event should be localised to nakkehoved is unclear. the profound changes in temperature over a short distance around nakkehoved are unlikely to be due to differences in heat flow but are more likely to reflect localised heating or differential exhumation (for further discussion, see section 4.3.1). håkansson et al. (1994) reported a vr value of c. 2% in a sample of the thyra ø formation from prinsesse margrethe ø, which indicates intense heating of this sample after deposition, specifically in post-paleocene times (håkansson et al. 1994). this observation, together with fig. 11 (continues) palaeotemperatures derived from afta. a: early permian (295– 290 ma). b: late triassic (225–210 ma). c: late jurassic (165–150 ma). d: paleocene (c. 60 ma). e: end-eocene (c. 35 ma). f: middle to late miocene (15–5 ma). palaeotemperature for the mid-cretaceous episode (115–90 ma) is not shown as only one sample falls within the map frame. the small arrow in the lower right corner of a, e and f indicates sample gc1113-48, which is located just south of the figure frame. geology legend in fig. 4. numbers next to afta samples denote sample numbers. hffz: harder fjord fault zone. kctz: kap canon thrust zone. kr: kap rigsdagen. pr.t.ø: prinsesse thyra ø. tlfz: trolle land fault zone. https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 25 of 43 www.geusbul let in.org the extreme vr values of upper cretaceous sediments around nakkehoved, led håkansson & pedersen (1982, 2001) to argue for a short-lived ‘post-paleocene thermal event’. however, this vr value was based on only a small number of measurements and was considered to be unreliable (paech & estrada 2018). vr values between 0.5 and 1.19% reported by pedersen et al. (2018), mainly from the southern part of kilen, indicate maximum palaeotemperatures of 83–155°c. just one value lies outside this range (2.57%; 210°c) for cretaceous formations that have been thrusted and folded in large open folds. these maturity data were interpreted in terms of large-scale folds with higher palaeotemperatures in anticlines, lower palaeotemperatures in synclines and an overall warming-northwards trend from the coastal zone of kilen to inland areas (pedersen et al. 2018). this suggests that the mesozoic parts of the wandel sea basin and kilen represent late cretaceous rift basins that subsequently underwent north–south compression (svennevig et al. 2016). pedersen et  al. (2018) modelled the pre-folding position of the strata and maturity data along a 9 km long river section, which cuts large-scale folds and thrusts in the southern part of kilen. this was possible due to the moderate maturity along the section (vr between 0.5 and 1.2%) and the consequent limited thermal alteration of, for example, dinocysts, which ensured a good biostratigraphic control. structural restoration along the section revealed k r o n p c h r i s t i a n l a n d kilen ø hovgaard n d 48 10°w 83 °n 82 °n 81 °n 80 °n 15°w20°w25°w30°w35°w40°w 15°w20°w25°w30°w35°w pr.t.ø pr.m.ø pr.d.ø depotbugt nakkehoved kctz hffz tlfz tlfz eg fzkk rrr oooo n pp cc h r i s t i a n l a n d kilen nnn dd t.øpr.t.ppr.t. m.ømpr.mm ør.d.øpr.dr.d.ø depotbugtdepotbudepotbu nakkehoved kctzzz hffzhhfff z tlfzlftlfz tlfz egegeg f g fzfz eg ffz paleocene–eocene sediments upper cretaceous sediments jurassic–cretaceous sediments equivalent palaeotemperature195°c 2.6% 2.6% 2.6% source of vr dataa, b, c number of samples(11) vr data paleocene–eocene sediments upper cretaceous sediments kap washington group jurassic–cretaceous sediments fault dykes, mostly upper cretaceous sedimentary outlier 3.2% (19) a 230°c 2.1% (7) a 195°c 1.3–2.9% (22) a 150–195°c 0.55% (12) a 90°c 1.6–2.0% (21) a 175–190°c 2.05% (1) c 90°c 0.4% (2) c 66°c 0.45% (1) c 76°c 0.36–0.51% (3) c 59–84°c 0.55% (31) a 90°c 0.5–2.26% (39) b 83–210°c 9.9% (1) c >250°c 7.1% (1) c >250°c 9.4% (1) c >250°c 7.4% (1) c >250°c 2.6–5.4% (3) a >200°c fig. 13 fig. 15 fig. 14 100 km500 fig. 12 published vitrinite reflectance (vr) values from mesozoic and palaeogene units of the wandel sea basin. coloured boxes: vr values and corresponding maximum palaeotemperatures (burnham & sweeny 1989). source of vr data indicated by letters a: paech & estrada (2018), b: pedersen et al. (2018), c: håkansson et al. (1994). when values are available from multiple sources, those from paech & estrada (2018) are shown. pr.d.ø: prinsesse dagmar ø. pr.m.ø: prinsesse margrethe ø. pr.t.ø: prinsesse thyra ø. detail of kilen in fig. 15. https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 26 of 43 www.geusbul let in.org a deeper pre-deformation burial of the northern area, which pedersen et al. (2018) interpreted as rift-related maximum burial after the santonian, which is the age of the youngest sediments on kilen (hovikoski et al. 2018). the subsequent folding and thrusting that occurred during basin inversion were assumed by pedersen et al. (2018) to be of the paleocene–eocene age. 4.3 comparison of afta and vr data in this study, we compare thermal history interpretations from afta and vr data in mesozoic and palaeogene units in the region from herluf trolle land to kronprins christian land (figs. 12–15). this area is chosen because variation in vr levels across the region, combined with the thermal history solutions from afta, provides key insights into the mechanisms responsible for the observed maturity levels. thermal history constraints for cenozoic palaeothermal episodes derived from afta and vr data are shown in figs. 13–15. these figures do not show the mesozoic cooling events identified from afta, as the vr data are measured in units which post-date these events, and because the afta results show that three cenozoic events dominate the thermal history of this region (fig. 10). 4.3.1 northern peary land in northern peary land, a small number of samples close to the hffz cooled below c. 110°c in the paleocene event (fig. 12). however, most samples in this region cooled below c.110°c in the end-eocene event. given that some samples appear to have cooled in both episodes, it seems likely that samples that cooled below c. 110°c at c. 35 ma also cooled at 60 ma, but remained at, or were reheated to palaeotemperatures >110°c until cooling that began at the end of the eocene. it is possible that cooling from palaeotemperatures of c. 200°c, represented by the higher vr values in the region between the kcfz and hffz, may have begun prior to 60 ma in a separate event, which cannot be resolved from the afta data. however, the simplest interpretation is that cooling from c. 200°c in northern peary land began in the regional paleocene episode at c. 60 ma. 4.3.2 herluf trolle land the mean vr value of 0.55% from paech & estrada (2018) in 12 lower cretaceous samples from southern herluf trolle land indicates a maximum palaeotemperature of c. 90°c (fig. 13). paech & estrada (2018) considered this level of maturity to be the regional ‘background’ level away from major fault systems and interpreted it to reflect the earlier presence of a considerable overburden across the region. afta data in samples gc1113-63 and -64, from this area (south-west of fault f in fig. 13), reveal end-eocene palaeotemperatures of 75–90°c and 70–80°c, respectively, from which cooling began at c. 35 ma. samples gc1113-21 and -22, also south of fault f, define similar end-eocene palaeotemperatures. these end-eocene palaeotemperatures from afta are broadly consistent with the maximum palaeotemperatures around 90°c, indicated by the vr data in nearby upper jurassic – lower cretaceous units (paech & estrada 2018). this consistency suggests that vr values reflect maximum paleotemperatures in the end-eocene episode. we, therefore, interpret maximum palaeotemperatures of 75–90°c, as defined from afta and vr in this region, to primarily reflect deeper burial prior to exhumation, which began at the end of the eocene. these values are relatively uniform over a wide area. in contrast, vr values up to 3% in the upper cretaceous herlufsholm strand formation from northern herluf trolle land indicate maximum palaeotemperatures at around 150–200°c within the prominent fault zone (fig. 13). afta data from this region provide consistent evidence of palaeotemperatures >100°c from which cooling began in the paleocene at c. 60 ma (gc1113-15, -59, -67 and -68). as all tracks were totally annealed prior to cooling that began at c. 60 ma, afta data in these samples define only lower limits to the maximum palaeotemperature, and direct comparison with the vr data is not possible. the most straightforward interpretation is that cooling from a palaeotemperature of c. 200°c, as indicated by the vr data, began at c. 60 ma. furthermore, mid-paleocene palaeotemperatures are strongly controlled by the faults of the tlfz (fig. 13; paech & estrada 2018). samples to the north-east of fault f cooled at c. 60 ma (gc1113-15, -23, -59, -67, -68), whereas samples to the south-west of that fault do not show any evidence of this event (gc1113-21, -22, -63, -64). afta data from the five samples north-east of fault f also show evidence of cooling beginning at 35 ma from palaeotemperatures of c. 80–90°c, similar to those defined from afta in the four southern samples, and thus, providing further evidence that this episode represents relatively uniform regional exhumation. 4.3.3 east of herluf trolle land the mean vr value of 0.55% in 31 palaeogene samples from prinsesse thyra ø corresponds to a palaeotemperature of c. 90°c (fig. 14). vr data in three samples of palaeogene age from prinsesse ingeborg halvø on kronprins christian land, as well as a lower cretaceous sample from kap rigsdagen, define maximum palaeotemperatures of 66–90°c. these are consistent with maximum post-depositional palaeotemperatures of 20–75°c and 80–90°c derived from afta data in samples gc1113-9 and -10, respectively. these palaeotemperatures are again attributed to the end-eocene palaeothermal episode and are consistent with the https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 27 of 43 www.geusbul let in.org values from herluf trolle land that are also attributed to deeper burial prior to regional exhumation, which began at c. 35 ma. along the north coast of kronprins christian land and in parts of kilen, vr reaches very high values in samples of cretaceous age: 1. around nakkehoved, vr values for upper cretaceous units of 7-10% indicate maximum palaeotemperatures in excess of 250°c (fig. 14). afta data in sample gc1113-55, also from nakkehoved, indicate that cooling from >125°c began between 101 and 48 ma. the simplest interpretation is that these values represent the paleocene (c. 60 ma) palaeothermal episode. 2. in 21 samples from inland kilen, vr values are in the range of 1.6-2% (equivalent to c. 174°c and 200°c, respectively; fig. 15). afta data in sample gc1113-58 from this area also define cooling from >120°c, which began between 60 and 45 ma, supporting that these elevated maturity levels reflect the palaeothermal episode at 60 ma. however, afta data in nearby sample gc1113-1 define cooling from >130°c between 45 and 25 ma, which we attribute to the cooling episode that began at the end of the eocene. it could be that this sample initially underwent paleocene cooling from a higher palaeothermal maximum, and then remained at >130°c until further cooling in the endeocene episode. in any event, the more recent episode was clearly more pronounced in sample gc1113-1 compared with sample gc111358. the higher end-eocene palaeotemperature quaternary cover u. cretaceous (herlufsholm strand fm) u. jurassic – l. cretaceous (ladegårdsåen fm) triassic (trolle land gp) permian (midnatsfjeld, kim fjelde fms) carboniferous (foldedal fm) silurian (peary land gp) proterozoic rocks ice fault sea and lake herlufsholm strand herlufsholm strand fs 70–80°c, 35–0 ma 75–80°c, 140–25 ma 15–75°c, 35–0 ma 85–95°c, 75–35 ma 25–50°c, 20–0 ma 85–100°c, 55–10 ma 22 21 63 64 afta sample without 60 ma event >100°c, 70–20 ma >105°c, 70–40 ma 55–50°c, 35–10 ma >130°c, 65–45 ma 95–105°c, 45–35 ma 30-65°c, 35–5 ma >120°c, 60–45 ma 76–103°c, 48–26 ma 25–72°c, 31–8 ma >100–105°c, 125–30 ma 20–95°c, 60–0 ma >100–105°c, 90–60 ma 68–85°c, 45–15 ma 67 59 5 68 23 65 afta sample with 60 ma event 1.3–2.9% (22) a 150–195°c 0.55% (12) a 90°c fault f herluf trolle land 20°w 82 °4 5’ n 82 °3 0’ n 21°w22°w23°w 5 km0 fig. 13 cenozoic palaeothermal constraints for herluf trolle land. heating during the paleocene event (c. 60 ma) is controlled by the numerous faults belonging to the trolle land fault zone, as suggested by paech & estrada (2018). only samples north-east of ‘fault f’ were affected by the paleocene thermal event. white boxes: thermal history solutions from afta data; colour coding as in fig. 10. coloured boxes: vr values and corresponding maximum palaeotemperatures from adjacent samples; see fig. 12. source of vr data indicated by letter a: paech & estrada (2018). map location indicated in fig. 12. basemap modified from bjerager et al. (2019). https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 28 of 43 www.geusbul let in.org in sample gc11113-1 emphasises the complex nature of the underlying thermal history and highlights the need for further studies in this region. 3. in more coastal areas of kilen, vr values for the cretaceous sediments are high and quite variable across the region (most values between 0.5 and 1.19%, equivalent to 83–155°c; fig. 15). the variation in vr values was attributed to differing amounts of burial prior to exhumation due to basin inversion of ‘presumed paleocene–eocene age’ (pedersen et al. 2018). the only direct comparison between afta and vr data in this region is sample gc1113-25, which cooled from >100°c between 65 and 10 ma. we attribute this to the paleocene palaeothermal episode from which cooling began at c. 60 ma. this is again consistent with the timing of cooling from extreme palaeotemperatures, which is indicated by the elevated vr levels further north in nakkehoved. given the considerable variation in vr values along the western margin of kilen and the lack of afta data from this region, it is not possible to draw any conclusions regarding the timing of cooling there. afta data in samples gc1113-8 and -60, from amdrup land, south of kilen, define maximum post-jurassic palaeotemperatures of c. 55–75°c (fig. 14), which is attributed to the end-eocene (c. 35 ma) episode. these palaeotemperatures are similar to those derived from afta and vr data obtained from areas to the north – away from areas of significant deformation – and provide further evidence of relatively uniform deeper burial across the region prior to the onset of exhumation at the eocene–oligocene boundary. 4.4 synthesis elevated vr values (1.5–3%; equivalent to 150–225°c) in mesozoic units within the major fault zones from northern peary land to kronprins christian land represent a mid-paleocene palaeothermal maximum from which d ø hovgaard p e a r y l a n d 48 83 °n 82 °n 81 °n 80 °n 15°w20°w 10°w15°w20°w 81 °n 82 °n tlfz d tlfz afta sample 40 km0 >120°c, 101–48 ma 86–106°c, 56–15 ma 36–81°c, 20–0 ma 55 20–75°c, 45–0 ma 80–90°c, 60–0 ma 910 8 60 45–75°c, 50–0 ma 55–75°c, 70–15 ma 2.05% (1) c 90°c 0.36–0.51% (3) c 59–84°c fig. 13 prinsesse thyra ø kap rigsdagen amdrup land nakkehoved fig. 15 0.4% (2) c 66°c 0.55% (31) a 90°c 7.1% (1) c >250°c fig. 14 cenozoic palaeothermal constraints east of herluf trolle land. white boxes: thermal history solutions from afta, colour coding as in fig. 10. coloured boxes: vr values and corresponding maximum palaeotemperatures from adjacent samples, colour coding as in fig. 12. source of vr data indicated by letters a: paech & estrada (2018), c: håkansson et al. (1994). solid black lines: faults. map location indicated in fig. 12. detail of kilen in fig. 15. key to bedrock geology and other abbreviations in fig. 4. https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 29 of 43 www.geusbul let in.org cooling began at c. 60 ma. in contrast, lower vr values (c. 0.55%, equivalent to c. 90°c), in mesozoic and palaeogene units located away from zones of major deformation, represent the end-eocene palaeothermal episode from which cooling began at c. 35 ma. afta data from the zones of deformation also show evidence of cooling, which began at this time from palaeotemperatures of c. 80–90°c, thus, emphasising the regional nature of  this episode, although end-eocene palaeotemperatures were higher in peary land north of the hffz (see section 5.2.2). a major exception to this regional pattern occurs on the north coast of kronprins christian land around nakkehoved. extremely high vr values between 7 and 10% in samples of upper cretaceous sediments indicate maximum palaeotemperatures greater than 250°c. afta data in the only sample from this region indicate that cooling from >125°c began in the paleocene or possibly earlier, since results provide only minimum palaeotemperature estimates. this event may, therefore, have been synchronous with paleocene cooling in the fault zones, or may have been earlier and, as the event affected upper cretaceous sediments, it must have happened in late cretaceous or paleocene times. afta data in sample gc1113-1 from kilen define cooling from >130°c during the end-eocene cooling episode, suggesting some variation in the magnitude of this episode. the vr value of c. 2% in a sample of palaeogene age on prinsesse margrethe ø may also reflect the higher end-eocene palaeotemperature in sample gc1113-1 (note the earlier comments regarding the uncertainty associated with this vr value). we suggest that exhumation that began in the mid-paleocene represents the inversion of the fault zones, while exhumation that began at the end of the eocene was regional and primarily represents burial followed by uplift and erosion however, there is also evidence of end-eocene palaeotemperatures that were 81°10’n 81°05’n 13°w13°30’w anduin gåseslette tågekyst sølverbæ k k a n g o q r y g sølverbæk : kr kr kr kr kr kr sb q kr tk? tk tk sb?sb sb sb 5 km hut and landing strip iskap sølverbæk fm, quarternary ice sea and lake undivided kangoq ryg mb (galadriel fjeld fm) tågekyst mb (galadriel fjeld fm) u pp er c re ta ce ou s sb kr tk lo w er c re ta ce ou s afta sample >120°c, 60–45 ma 76–103°c, 48–26 ma 25–72°c, 31–8 ma >130°c, 45–25 ma >100°c, 65–10 ma 58 25 1 0.66% b 109°c 2.57% b 210°c 0.72% b 118°c 0.92% b 137°c 1.04% b 145°c 0.96–1.05% b 140–145°c 0.78–1.19% b 125–155°c 0.55–0.70% b 91–116°c 0.50–0.60% b 83–99°c 0.58% c 95°c fig. 15 palaeothermal constraints for south-east kilen. the palaeotemperature constraints from vr are consistent with those from afta, but the timing constraints on the onset of cooling vary. white boxes: thermal history solutions from afta, colour coding as in fig. 10. numbers next to afta samples denote sample numbers. blue boxes: vr values and associated maximum palaeotemperatures for cretaceous sediments. source of vr data indicated by letters b: pedersen et al. (2018) and c: håkansson et al. (1994). map location in fig. 14. base map modified from hovikoski et al. (2018, fig. 4) where further details of the map are explained. https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 30 of 43 www.geusbul let in.org higher than the regional, uniform magnitude, for example, sample gc1113-1 from kilen and the possibly highly mature thyra ø formation on prinsesse margrethe ø. the extreme palaeotemperatures focussed along the north coast of kronprins christian land were likely reached during the late cretaceous or paleocene times. 5 geological controls on the cooling episodes identified from afta in this section, we compare afta and vr results with evidence from the stratigraphic record to investigate the geological processes responsible for the cooling episodes identified in the wandel sea basin (fig. 16). we also compare these cooling episodes with studies from central west greenland and north-east greenland (fig. 17; japsen et al. 2006, in press) and illustrate the heating and cooling (burial and exhumation) history, with examples based on samples that provide stratigraphic constraints on the age of removed sections (fig. 18). 5.1.palaeozoic–mesozoic cooling events 5.1.1 early permian episode the early permian cooling episode (beginning between 295 and 290 ma) recognised from afta data in samples across the study area corresponds to the base-permian hiatus identified across the basin between the middle– upper permian kim fjelde formation and the upper carboniferous foldedal formation (fig. 16; stemmerik et al. 1998). we, therefore, interpret early permian cooling as reflecting a major phase of exhumation, following deeper burial. for example, sample gc1113-63, a silurian sandstone from herluf trolle land, cooled from an early permian palaeotemperature of c. 100°c, corresponding to burial below a cover of silurian to carboniferous rocks with a thickness of 2.7 km (30°c/km, 20°c at the surface). the timing of this cooling episode is close to that of the late carboniferous episode identified in north-east greenland (beginning between 320 and 300 ma), which also correlates with a regional latest carboniferous to mid-permian unconformity (surlyk 1990). it is, therefore, possible that these episodes reflect the same tectonic event (fig. 17). it is not possible to define any pre-permian cooling from the afta data, and therefore, it is not possible to use afta data to discriminate the thermal histories for upper and lower carboniferous strata on holm land based on maturity data, as in stemmerik et al. (1998). 5.1.2 late triassic episode the late triassic cooling episode (beginning between 225 and 210 ma) is identified from afta data away from the northern coastal regions. the onset of cooling in this episode could correspond to the older hiatus between the upper triassic sortekløft formation and the middle triassic dunken formation, identified on herluf trolle land by bjerager et al. (2019; fig. 16). in kilen, the episode corresponds to the inferred hiatus between the middle jurassic mågensfjeld formation and the middle triassic isrand formation (alsen et al. 2017; hovikoski et al. 2018; svennevig et al. 2018). sample gc111318 of upper carboniferous kap jungersen formation exposed on holm land reached a palaeotemperature above 100°c prior to late triassic cooling (fig. 18a). this palaeotemperature corresponds to burial below a cover of carboniferous to triassic sediments, with a thickness exceeding 2.7 km (30°c/km, 20°c at the surface). the late triassic episode in the wandel sea basin coincides with the late triassic episode in west greenland, but slightly post-dates a middle triassic episode of cooling and exhumation in north-east greenland (fig. 17), suggesting a possible lag between events in the two areas. 5.1.3 late jurassic episode the late jurassic cooling episode (beginning between 165 and 150 ma) is recognised in samples away from the northern coasts and in one sample from herluf trolle land. the onset of cooling in this episode corresponds to the hiatus between the upper jurassic – lower cretaceous ladegårdsåen formation and the upper triassic sortekløft formation on herluf trolle land (håkansson & pedersen 2015; bjerager et al. 2019; fig. 16). we interpret this cooling episode as exhumation following deeper burial. according to håkansson & pedersen (2015), the base of ladegårdsåen formation is middle oxfordian in age (c. 160 ma), which constrains the onset of late jurassic cooling to between 165 and 160 ma, at the middle–late jurassic transition (callovian– oxfordian). this agrees well with the callovian–oxfordian hiatus between the birkelund fjeld and mågensfjeld formations observed in kilen (hovikoski et al. 2018). the temporal overlap between the late jurassic episode and the callovian–oxfordian hiatus across most of the basin indicates that the cooling episodes involved exhumation. for example, sample gc1113-22 from herluf trolle land of lower triassic sandstone reached a palaeotemperature above 115°c prior to late jurassic cooling (fig. 18b). this corresponds to burial below a cover of triassic to middle jurassic sediments, with a thickness exceeding 3 km (30°c/km, 20°c at the surface). the late jurassic episode that affected the wandel sea basin coincides with the late jurassic episode that affected west greenland, but it began after the early jurassic episode of cooling and exhumation in north-east greenland (fig. 17). as this delay of about 20 myr is consistent with the delay of the triassic episode in north-east greenland compared with western and northern greenland, it seems likely to represent https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 31 of 43 www.geusbul let in.org is ra nd kap jungersenkap jungersen pb sortebakker mågensfjeldmågensfjeld herlufsholm strand la de gå rd så en sølverbæk ??? ??? thyra ø kw nakkehoved ld d lfoldedal kim fjelde lichenrygyg endunken storekløft elddemidnatsfje g al ad ri el f je ld bf dd klklkl peary land k.c. land wandel sea basin j channel fills marine mudstone marine carbonate marine sandstone sedimentary environments horizontal shading lithological symbols coal layers sandstone siltstone mudstone/shale organic rich mudstone/shale limestone fluvial, lacustrine extrusive igneous rocks volcanics, volcanoclastics evaporites conglomerate legend time interval during which cooling began ma era stage/ age eocene oligocene miocene pliocene pleistocene series/ epoch system/ period middle middle lower upper guadalupian lopingian lower lower middle mississippian upper mississippian lower penns. middle penns. upper penns. cisuralian lower mississippian langhian aquitanian chattian rupelian priabonian bartonian lutetian thanetian selandian danian maastrichtian campanian santonian coniacian aptian hauterivian valanginian berriasian tithonian bathonian bajocian aalenian pliensbachian sinemurian hettangian carnian ladinian anisian olenekian induan rhaetian roadian wordian capitanian wuchiapingian changhsingian kungurian artinskian asselian gzhelian kasimovian moscovian bashkirian serpukhovian visean tournaisian barremian toarcian ypresian burdigalian pa la eo zo ic m es o zo ic c en o zo ic n eo g en e pa la eo g en e c re ta c eo u s ju ra ss ic tr ia ss ic pe rm ia n c a rb o n if er o u s sakmarian upper kimmeridgian oxfordian albian cenomanian turonian serravallianserravallian tortonian messinian upper norian paleocene callovian 0 10 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 210 220 230 240 250 260 270 280 290 300 310 320 330 340 350 fig. 16 comparison between the regional cooling episodes identified from afta (table 4) and the stratigraphic scheme for the wandel sea basin (fig. 5). the thickness of the horizontal, coloured bars indicates the interval during which the cooling began in each episode. a protracted tectonic development between the two areas. 5.1.4 mid-cretaceous episode the mid-cretaceous cooling episode (beginning between 115 and 90 ma) is identified in only one sample within the wandel sea basin. the episode corresponds to an aptian to turonian hiatus on herluf trolle land between the upper cretaceous (turonian–coniacian) herlufsholm strand formation and the upper jurassic – lower cretaceous (oxfordian–barremian) ladegårdsåen formation (fig. 16) (håkansson & pedersen 2015; piasecki et al. 2018). in kilen, hovikoski et al. (2018) marked a possible late albian to cenomanian hiatus between the galadriel fjeld and sølverbæk formations in the kilen fjelde area. the temporal overlap between this hiatus and the mid-cretaceous cooling episode suggests that cooling involved exhumation following deeper burial. the mid-cretaceous episode of uplift and erosion postdates the barremian unconformity (c. 135 ma), which resulted from halip-related regional uplift and is recorded by stratigraphic and palaeo-environmental changes on peary land and on kilen (ineson et al. 2020). it is possible that the effects of the barremian episode overlaps with those of the mid-cretaceous event as the https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 32 of 43 www.geusbul let in.org fig. 17 timing of regional, post-devonian episodes of uplift and erosion in three regions of greenland estimated from afta. a: central west greenland (japsen et al. 2006, 2009). b: eastern north greenland (table 4; this study). c: north-east greenland (japsen et al. in press). break-up west of greenland took place in mid-paleocene (c. 62 ma), but sea-floor spreading ceased there by the end of the eocene. in contrast, the opening of the north-east atlantic, east of greenland, began at the paleocene–eocene transition (c. 56 ma) and is still ongoing (chalmers & pulvertaft 2001; oakey & chalmers 2012; gaina et al. 2017). the fram strait between greenland and svalbard opened in early miocene (c. 20 ma; jokat et al. 2016). l: late. e: early. m: middle. c: carboniferous. pe: permian. tr: triassic. j: jurassic. kl: lower cretaceous. ku: upper cretaceous. pal: paleocene. eo: eocene. oli: oligocene. mio: miocene. pl: pliocene. q: quaternary. 30 0 20 0 10 0 65 60 40 20 0 30 0 20 0 10 0 65 60 40 20 0 central west greenland north-east greenland 30 0 20 0 10 0 65 60 40 20 0 eastern north greenlanda b c m.–l. miocene (15–5) end–eocene (c. 35) l. jurassic (165–150) l. triassic (225–210) e. permian (295–290) paleocene (c. 60) m. cretaceous (115–90) break-up q pl mio oli eo pal ku kl j tr pe c q pl mio oli eo pal ku kl j tr pe c q pl mio oli eo pal ku kl j tr pe c time interval during which cooling began (7–2) onset of cooling (ma) from afta ti m e (m a) eocene–oligocene (36–30) (reburial) l. miocene (11–10) l. neogene (7–2) break-up l. jurassic (160–150) l. triassic (230–220) e. pliocene (c. 5) l. miocene (c. 10) end–eocene (37–35) break-up e. jurassic (c. 180) m. trassic (c. 240) e. cretaceous (145–140) m. cretaceous (95–90) l. carboniferous (320–300) one sample within the wandel sea basin, which has been assigned to the latter event cooled between 155 and 50 ma. the onset of cooling in this episode coincides with a mid-cretaceous episode in north-east greenland (fig.  17), which represents an episode of regional exhumation (japsen et al. in press). we, therefore, suggest that both areas were affected by exhumation at 95–90 ma. 5.2 cenozoic cooling episodes 5.2.1 mid-paleocene episode the mid-paleocene cooling episode (beginning c. 60 ma) recorded in samples from coastal and near-coastal locations corresponds to the hiatus between the upper paleocene – lower eocene thyra ø formation and upper cretaceous strata (e.g. santonian sølverbæk formation in kilen; fig. 16). the study results from the hffz and the tlfz reveal paleocene cooling from >110°c, which began at c. 60 ma. we interpret this cooling in terms of exhumation following deeper burial (fig. 11d). this event is also inferred to have affected northern peary land in the region between the kctz and hffz where high end-eocene palaeotemperatures dominate. as discussed in section 4.3.3, afta and vr data from kilen and northern herluf trolle land indicate that cooling from a maximum post-depositional palaeotemperature of up to 200°c within major fault zones also began during the paleocene (c. 60 ma) palaeothermal episode. the results of this study support the interpretation of pedersen et al. (2018), according to which the high maturity of sediments in the southern part of kilen was reached during the maximum burial prior to post-coniacian basin inversion in response to a strong compressional event. for example, sample gc1113-58 of lower cretaceous lichenryg formation exposed in kilen reached a palaeotemperature of >120°c prior to paleocene cooling (fig. 18c). this corresponds to burial beneath more than 3.3 km of cretaceous to paleocene sediments. deep burial of the cretaceous sediments in kilen is in good agreement with the occurrence of well-developed stylolites in the lower part of lichenryg formation in the northern part of kilen (around dromledomen; c. heinberg & h. dypvik, personal communication 2019). similar considerations apply to https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 33 of 43 www.geusbul let in.org samples gc1113-5 and -59 of the upper cretaceous herlufsholm strand formation (herluf trolle land) that reached palaeotemperatures >100°c prior to paleocene cooling. in summary, we observe a spatial and temporal overlap between the inversion of the major fault zones of the wandel sea basin and the onset of cooling and exhumation that began at c. 60 ma. we infer that the compression that resulted in inversion of the fault zones began in the mid-paleocene (i.e. the kronprins christian land orogeny; håkansson & pedersen 1982; svennevig et al. 2016). this timing is consistent with the late paleocene age of the lower part of the thyra ø formation, which may represent erosional products derived from the inversion of the fault zones. this interpretation is supported by the abundant of reworked cretaceous dinoflagellate cysts and palynomorphs in these deposits (lyck & stemmerik 2000; piasecki et al. 2018). a corresponding episode is not revealed by afta data from north-east greenland or west greenland (fig. 17). however, the timing of this episode overlaps with that of a late cretaceous – early cenozoic episode (beginning between 70 and 50 ma), which affected the southernmost tip of east greenland (63°–61°n; green et al. 2014). dam et al. (1998) reported geological evidence for rapid uplift and fluvial erosion in the mid-paleocene followed by subsidence contemporaneous with the first volcanic extrusions in both west and east greenland, and argued that the arrival of the iceland plume controlled these transient, vertical movements. 5.2.2 end-eocene episode the cooling episode that began at the end of the eocene (c. 35 ma) is recognised in afta data in samples across almost the entire study area. the cooling episode affected the wandel sea basin after the deposition of the thyra ø formation and coincides with periods of exhumation in north-east greenland, as well as west greenland (fig. 17). afta and vr data define fairly uniform palaeotemperatures of around 80–90°c away from major fault zones, corresponding to the regional (or background) level of maturity across the region. this is interpreted to represent primarily deeper burial prior to regional, relatively uniform exhumation that began at the end of the eocene. the palaeogene sediments on prinsesse thyra ø (e.g. gc1113-9) also reached such palaeotemperatures prior to the end-eocene episode, implying that the heating there reflects burial below a late paleocene to eocene cover. piasecki et al. (2018) showed that the top 5 m of the section exposed at kap rigsdagen are of early–mid eocene age and are part of the thyra ø formation. the end-eocene palaeotemperatures (80–90°c) for sample gc1113-10 from lower cretaceous sediments overlain by palaeogene sediments at kap rigsdagen must, therefore, reflect burial below a cover of eocene sediments of similar thickness to that across prinsesse thyra ø (fig.  18d). assuming a palaeogeothermal gradient of 30°c/km and a palaeosurface temperature of 10°c, this range of palaeotemperatures corresponds to burial below 2.3–2.7 km of eocene sediments. the choice of palaeosurface temperature used here reflects the warm temperate climate that boyd (1990) estimated based on the palaeogene flora preserved in the deposits on prinsesse thyra ø. it is thus likely that about 2.5 km of upper paleocene to eocene sediments (including the thyra ø formation) covered wide areas across the wandel sea basin prior to the exhumation that began at the end of the eocene. this suggests that large parts of the basin subsided during the eocene. some samples from north of the hffz (e.g. gc1113-15, -43, -45) cooled from much higher palaeotemperatures (>110°c) compared with the majority of samples across the region (see section 6.3). afta data for sample gc1113-1 from the northern part of kilen also indicate cooling from >130°c between 45 and 25 ma during the end-eocene episode. 5.2.3 mid-late miocene episode the mid–late miocene cooling episode (beginning between 15 and 5 ma) is identified in the afta data from samples in coastal and near-coastal locations, and does not correlate with any known strata within the wandel sea basin. it does, however, correlate with the late miocene and early pliocene events of uplift and incision that led to the formation of the present-day landscape in both west and north-east greenland (fig. 17; japsen et al. 2006, in press). as the palaeotemperatures defined for this episode are moderate (c. 50–60°c at sea level), they are unlikely to be related to the early pliocene cooling observed in north-east greenland (japsen et al. in press). we thus suggest that the mid-late miocene episode identified in this study shares a common origin with the late miocene episode (c. 10 ma) in north-east greenland, reflecting cooling due to uplift and erosion of the continental margin and development of the modern-day topography. afta data from herluf trolle land consistently define miocene palaeotemperatures around 50°c at sea level, corresponding to burial below a 1.5 km thick rock column (30°c/km, 5°c at the surface). miocene palaeotemperatures in samples from northern peary land and from kronprins christian land are less well constrained but are nonetheless consistent with a similar value. the failure to resolve the miocene event at locations away from the northern coasts may indicate that miocene https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 34 of 43 www.geusbul let in.org palaeotemperatures were lower than, for example, on herluf trolle land. figure 19 shows a three-dimensional elevation model of the region around independence fjord and illustrates that independence fjord is incised below a low-relief surface at just above 1 km a.s.l., in contrast to the more elevated and dissected alpine relief further north on peary land. incision below that low-relief surface would result in cooling of rocks now near the sea level by <50°c, suggesting that the miocene palaeotemperatures may reflect cooling due to the incision of the present-day relief following miocene uplift. this contradicts the suggestion that neogene uplift and erosion of the wandel sea basin were very limited ( stemmerik et al. 1998). 6 discussion the climax of the eurekan orogeny is agreed to have occurred in the eocene (steel et al. 1985; piepjohn et al. 2016), but both the onset and the end of the orogeny have been a matter of debate. here, we apply afta data to place tighter constraints on the timing of these events. we also address the origin of the extreme heat recorded at nakkehoved on the northern coast of kronprins christian land and comment on the likely former extent of the wandel sea basin. 6.1 mid-paleocene onset of the eurekan orogeny we have demonstrated that cooling and exhumation along the hffz and the tlfz (including kilen) began time (ma) carboniferous te m pe ra tu re (° c ) devon eocenepaleocenecretaceousjurassictriassicpermian oligocene miocene p q deposition of carboniferous sediments a b present-day exposure ? ? 80–90ºc 80–110ºc >100ºc 0 20 60 80 100 120 40 400 300 20 10 0304050100 606520020 carboniferous te m pe ra tu re (° c ) devon eocenepaleocenecretaceousjurassictriassicpermian oligocene miocene p q deposition of lower triassic sediments present-day exposure ? ? 85–95ºc >115ºc 25–50ºc time (ma) 0 20 60 80 100 120 40 400 300 20 10 0304050100 6065200 heating below triassic to middle jurassic sediments heating below carboniferous to triassic sediments fig. 18 thermal history diagrams illustrating the heating and cooling (burial and exhumation) of four samples that provide stratigraphic constraints on the age of removed sections. a: gc1113-18 (20 m a.s.l.) of upper carboniferous kap jungersen fm exposed on holm land. the sample reached a palaeotemperature above 100°c prior to late triassic cooling. this palaeotemperature corresponds to burial below a cover of carboniferous to triassic sediments, with a thickness exceeding 2.7 km. b: gc1113-22 (200 m a.s.l.) of lower triassic parish bjerg fm exposed on herluf trolle land. the sample reached a palaeotemperature above 115°c prior to late jurassic cooling, corresponding to burial below a cover of triassic to middle jurassic sediments with a thickness exceeding 3.1 km.the conversion of palaeotecmperatures to burial depths is based on an assumed palaeogeothermal gradient of 30°c/km and palaeosurface temperatures of 10°c for the end-eocene episode and 20°c for earlier episodes. red triangles: constraints for conditions at the surface during deposition of sediments or at the present day. coloured boxes: constraints from afta on palaeotemperature prior to cooling (the width of the boxes is fixed for better overview). boxes extend below the x-axis of the diagram where only a lower limit of the palaeotemperature is known (e.g. >110°c). sample locations in fig. 6. https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 35 of 43 www.geusbul let in.org during the mid-paleocene episode (c. 60 ma), and a similar timing is likely for the region between the kcfz and hffz, northern peary land. this episode, therefore, likely defines the onset of compression along these fault zones. we follow the interpretation of pedersen et al. (2018) that the high palaeotemperatures in the southern part of kilen were reached during maximum burial, prior to basin inversion, during a strong compressional event that we thus date as mid-paleocene in age. the paleocene timing of the compressional event is consistent with the assertion of håkansson & pedersen (2015 and earlier studies) that the compressional tectonics in the wandel sea basin, the kronprins christian land orogeny, occurred around the cretaceous–palaeogene boundary. the onset of compression in the wandel sea basin coincides with the mid-paleocene onset of sea-floor spreading in the labrador sea and thus with the start of movement of greenland towards the east-northeast (oakey & chalmers 2012). this implies that the mid-paleocene compression of the wandel sea basin took place during the first stage of the eurekan orogeny. the relationship between the undeformed palaeogene strata and the heavily deformed upper cretaceous strata in the fault zones in the wandel sea basin is not known from outcrops. it is thus possible that the deformation continued in the fault zones during the deposition of the palaeogene sediments. it has been suggested d c carboniferous te m pe ra tu re (° c ) devon eocenepaleocenecretaceousjurassictriassicpermian oligocene miocene p q deposition of lichenryg fm present-day exposure ?? 85–105ºc >120ºc 35–70ºc 6060 time (ma) 0 20 60 80 100 120 40 400 300 20 10 0304050100 65200 carboniferous te m pe ra tu re (° c ) devon eocenepaleocenecretaceousjurassictriassicpermian oligocene miocene p q deposition of ladegårdsåen fm present-day exposure ? 80–90ºc thyra ø fm resting on ladegårdsåen fm time (ma) 0 20 60 80 100 120 40 400 300 20 10 0304050100 6065200 heating below eocene sediments heating below cretaceous to paleocene sediments fig. 18 (continued) c: gc1113-58 (264 m a.s.l.) of lower cretaceous lichenryg fm exposed in kilen. the sample reached a palaeotemperature above 120°c prior to paleocene cooling, corresponding to burial below a cover exceeding a thickness of 3.3 km of cretaceous to paleocene sediments. d: gc1113-10 (23 m a.s.l.) of lower cretaceous sediments of the ladegårdsåen fm exposed at kap rigsdagen below a thin cover of the prinsesse thyra ø formation of early to middle eocene age (piasecki et al. 2018). the palaeotemperature of 80–90°c for the lower cretaceous sediments prior to end-eocene exhumation reflects heating below a thick cover of eocene sediments at this location. this range of palaeotemperatures corresponds to burial below 2.3–2.7 km of eocene sediments. the conversion of palaeotecmperatures to burial depths is based on an assumed palaeogeothermal gradient of 30°c/km and palaeosurface temperatures of 10°c for the end-eocene episode and 20°c for earlier episodes. red triangles: constraints for conditions at the surface during deposition of sediments or at the present day. coloured boxes: constraints from afta on palaeotemperature prior to cooling (the width of the boxes is fixed for better overview). boxes extend below the x-axis of the diagram where only a lower limit of the palaeotemperature is known (e.g. >110°c). sample locations in fig. 6. https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 36 of 43 www.geusbul let in.org that deformation at kap washington in the western part of the wandel sea basin continued into the eocene, as the ar-ar system of the kap washington group was partially reset in the early eocene (c. 47 ma; tegner et al. 2011). however, the ar-ar age spectra do not show well-defined plateaus, and the resulting ages should perhaps be treated with caution. jones et al. (2017) used u-pb zircon dates from tephra layers close to the base-paleocene unconformity of the ctb on the island of spitsbergen (in the svalbard archipelago) to define the onset of basin formation there as mid-paleocene, c. 61.8 ma. jones et al. (2017) also linked the formation of the ctb to the initiation of compression between eastern north greenland and svalbard, and they argued that the cessation of the kap washington volcanism around that time was corroborating evidence. had there been extension between these regions, then the activity would likely have increased (jones et al. 2017; fig. 5). thus, the formation of the ctb on spitsbergen began at about the same time as the compression in the wandel sea basin, slightly prior to 60 ma and prior to the deposition of the late paleocene to earliest eocene thyra ø formation. interestingly, the ctb only received sediments from only eastern source areas during the paleocene, and the basin received sediments only from the west after the earliest eocene emergence of the west spitsbergen fold belt (dallmann et al. 2015; petersen et al. 2016). the mid-paleocene timing for the onset of the compressional event is comparable with the late paleocene – early eocene timing suggested by døssing et al. (2010) and svennevig et al. (2016) but differs from the eocene timing suggested by von gosen & piepjohn (2003). however, according to von gosen & piepjohn (2003) and svennevig et al. (2016), a late paleocene or even eocene north–south compression of the wandel sea basin would agree with a unified timing of eurekan deformation extending from the canadian arctic archipelago to north greenland and svalbard. however, the eurekan deformation occurred in two stages, as discussed in section 2.2. during the first (paleocene) eurekan stage, the south-south-west to north-north-east movement of greenland relative to svalbard gave rise to compression that caused exhumation in the wandel sea basin and subsidence of the ctb during its formation as a foreland basin. during the second (eocene) eurekan stage, compression continued, but the greenland plate changed trajectory resulting in sinistral strike-slip transpression between greenland and svalbard. much of what can be seen today in the field is the final result of north–south transpression during the second eurekan stage. independent lines of evidence support a mid-paleocene onset of the eurekan orogeny. the sverdrup basin in the north-eastern canadian arctic archipelago underwent a drastic change in the earliest paleocene when the basin began to deform, coincident with the onset of sea-floor spreading west of greenland (embry & beauchamp 2019). using apatite fission-track data, arne et al. (2002) reported significant paleocene cooling in the sverdrup basin, and green & duddy (2010) reported the onset of extensive cooling in the paleocene between 60 and 55 ma in alaska, arctic canada, the northern barents sea and svalbard based on afta data. the plate reconstructions published by müller et al. (2016) showed that transpression with 18° convergence between the barents margin and north greenland was initiated in the late paleocene (60 ma) and culminated in the early eocene (54–50 ma; svennevig et al. 2016). we conclude that the mid-paleocene episode of cooling and exhumation (c. 60 ma) along the major fault zones in the wandel sea basin reflects the onset of the eurekan orogeny due to the convergence between greenland and the barents margin. 6.2 recent results from arctic canada recent apatite fission-track and apatite (u-th-sm)/he data from a suite of samples from the pearya terrane in northern ellesmere island (analogous to area 1 of this study; fig. 6) yielded fission-track ages of c. 40–50 ma over an elevation range of c. 1600 m (vamvaka et  al. 2019). he-ages in individual grains showed a much wider range, from which the authors selected certain values to derive thermal history constraints. the logic behind this selection is not clear, as the ages show no relationship with either equivalent u content or grain size, as required by the models used to interpret these data (see green & duddy [2012] for discussion). from the combination of afta and he data, they presented a preferred scenario with three periods of cooling, showing a close correspondence with the three episodes defined on independent grounds by piepjohn et al. (2016). this differs significantly from the preferred thermal history synthesis for wandel sea basin presented here, in which major cooling occurred at c. 60 and 35 ma. however, we noted from the supplementary information presented by vamvaka et al. (2019) that their solutions were constrained to begin cooling from temperatures of 140– 160°c between 60 and 50 ma, which is consistent with the paleocene cooling episode defined from afta in this study. it, therefore, seems likely that paleocene cooling in the two areas has a common, tectonic origin, although vamvaka et al. (2019) considered it to have a non-tectonic origin. on comparing our data with those of vamvaka, it seems unlikely that their data could independently define three cooling episodes between 55 and 35 ma. instead, we believe that a more reasonable interpretation would be two major cooling episodes, which began https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 37 of 43 www.geusbul let in.org about 60 and 35 ma, as in north greenland, suggesting a similar tectonic evolution in both north greenland and ellesmere island. 6.3 end-eocene tectonics in relation to the eurekan orogeny there is a clear offset in end-eocene palaeotemperatures across the hffz (fig. 11e). typical palaeotemperatures correspond to burial below a cover of c. 3.5 km and c. 2.5 km a.s.l. on the northern and the southern side of the fault, respectively (c. 110°c and 85°c at sea level; 30°c/km, 10°c at the surface). this implies that the rocks on the northern side of the fault were more deeply buried by about 1 km, compared with those south of the fault, prior to end-eocene exhumation. the present-day summit level on both sides of the hffz is about 1.8 km a.s.l., corresponding broadly to the thickness of the cover needed to explain miocene palaeotemperatures of c. 60°c at sea level on both sides of the fault, as defined by afta from the high mountains north of the fault zone and from herluf trolle land (section 5.2.3). based on the landscape development in north-east greenland, we suggest that the summit level in this study area of up to 2 km a.s.l. represents a peneplain that was graded to near sea level after end-eocene exhumation and uplifted to the level of the present-day summits after c. 10 ma (bonow & japsen 2021; japsen et al. in press). the present-day relief was, therefore, likely shaped by differential uplift and subsequent incision below this peneplain since the late miocene. the digital elevation model shown in fig. 19 illustrates the variation of the relief across south-eastern peary land to the south of the hffz. the constant summit level across the hffz implies that there was no significant vertical movement along the fault during the mid–late miocene tectonic episode, and thus, that the offset of end-eocene palaeotemperatures is due to differential vertical movements across the hffz that had taken place by the end of the eocene. north–south compression affected northern peary land in post-santonian times, with reverse faulting at the hffz and thrusting along the kctz, which piepjohn & von gosen (2001) attributed to the eurekan orogeny. however, they also argued that compression along the hffz was terminated by the change in plate-tectonic configuration at c13 time (c. 35 ma), when greenland became part of the north american plate. our afta data document kilometre-scale, reverse movements along the hffz, which is consistent with the north–south reverse and thrust faults observed by piepjohn & von gosen (2001). however, the results of this study also show that the inversion of the hffz took place in two episodes that began (1) during the mid-paleocene (c. 60 ma) and (2) at the end of the eocene (c. 35 ma). the first episode represents the onset of the eurekan orogeny, while the second episode represents post-eurekan tectonics as the greenland plate had become part of the north american plate. consequently, the eurekan orogeny had terminated by the end of the eocene (as supported by piepjohn & von gosen 2001), and the reverse faulting along the hffz that occurred at that time is, therefore, unrelated to the eurekan orogeny. while the end-eocene cooling and exhumation began after the termination of spreading in the labrador sea, and hence after the eurekan orogeny, it also coincides with an important phase of plate reorganisation in north-east atlantic (gaina et al. 2009, 2017). the episode of uplift and erosion in north greenland that began at the end of the eocene is synchronous with episodes of exhumation in east and west greenland, far from the eurekan orogen (section 5.2.2). many other observations show that significant tectonic changes affected greenland and adjacent arctic regions around the eocene–oligocene transition. for example, late eocene magmatic activity affected the east greenland margin (larsen et al. 2014). uplift of the inner margin of south-east greenland resulted in a strong flux of coarse clastic turbidites during the late oligocene on the shelf, above a middle eocene – upper oligocene hiatus (larsen et al. 1994). end-eocene tectonic uplift led to the formation of regional peneplains on both sides of greenland during the oligocene and miocene time (japsen et al. in press). end-eocene exhumation affected arctic regions beyond greenland, including the sverdrup basin, canada, as well as the barents sea, svalbard and the north slope of alaska (green & duddy 2010; embry & beauchamp 2019). all these observations reveal that a deep-rooted, tectonic process began in the north-east atlantic at the end of the eocene after the eurekan orogeny. further studies, including geodynamic modelling, are needed to investigate the origin of these post-eurekan, tectonic changes. 6.4 the extreme heating recorded at nakkehoved vr data from around nakkehoved, on the north coast of kronprins christian land, document an extreme degree of heating of upper cretaceous sediments to palaeotemperatures >250°c within a very limited area (fig. 12). this extreme heating stands in stark contrast to much lower values (80–90°c) in surrounding areas, and to date, no convincing explanation has been offered for these high temperatures. nonetheless, they are thought to have occurred post-paleocene (håkansson & pedersen 2001, https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 38 of 43 www.geusbul let in.org 2015; paech & estrada 2018). for example, håkansson & pedersen (2001) differentiated between (1) the compressional tectonics of the kronprins christian land orogeny that culminated at the cretaceous– palaeogene boundary (corresponding to the mid-paleocene episode of exhumation defined here) and (2) a post-paleocene thermal event. these authors argued that the latter event was responsible for the elevated vr value of the tectonically undeformed sediments of the thyra ø formation on prinsesse margrethe ø, as well as for the extreme vr values around nakkehoved. afta data from most samples analysed from around nakkehoved and kilen show evidence of paleocene cooling from elevated temperatures. however, whether the palaeotemperatures on kilen were as high as those responsible for the vr levels of up to 7% or more around nakkehoved is not certain. the extreme palaeotemperatures may have been reached before 60 ma, as afta data from nakkehoved and kilen provide only minimum palaeotemperature estimates for this time. however, as the event affected upper cretaceous sediments, it may thus have happened in late cretaceous or paleocene times. given the complexity of the region around nakkehoved, additional sampling for afta and vr, combined with detailed stratigraphic and sedimentological studies is required before the cause of the extreme thermal imprint there can be understood. 6.5 previous extent of the wandel sea basin substantial thicknesses of sediments of carboniferous to palaeogene age have been present within the present-day outline of the wandel sea basin and outside the relatively narrow late cretaceous rift zones (fig. 20). the kilometre-scale thickness of these now removed sedimentary covers implies that they must have extended substantial distances beyond the present-day outline of the wandel sea basin. when we consider the reconstructed geographical positions of carboniferous–palaeogene basins in north-east greenland, svalbard and the barents shelf prior to cenozoic sea-floor spreading (gion et al. 2017), it is likely that the wandel sea basin was at times coherent with them. several authors have highlighted the marked similarities among the sedimentary sequences in these three regions (håkansson & stemmerik 1989; mørk et al. 1989; von gosen & piepjohn 2003). håkansson & stemmerik (1984) presented evidence for a parallel development of the wandel sea basin with svalbard and the barents shelf. they found that the depositional environments of the wandel sea basin in the carboniferous to triassic and in the paleocene compared well with known stratigraphy in svalbard, whereas the late mesozoic history of the two areas showed major differences. in particular, the presence of thick upper cretaceous sequences in north greenland constitutes the most conspicuous n 0 200 400 600 800 1000 1200 1400 1600 1800 2000 2200 2400 m independence fjord wandel dal fig. 19 3d-elevation model of the region around independence fjord (globe 1999; map location in fig. 3). note that independence fjord and wandel dal (valley in danish) are incised below a low-relief surface at about 1 km a.s.l. this is in contrast to more elevated and dissected, alpine relief north-west of independence fjord. the width of the diagram is about 250 km. illustration: johan m. bonow. https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 39 of 43 www.geusbul let in.org difference from svalbard, where sediments of that age are absent (dallmann et al. 2015). a study of the late palaeozoic in the north atlantic found evidence for a change in the overall depositional environment of the north greenland – barents sea area from extensive, humid flood plains in the early carboniferous, to shallow warm seas in the mid-carboniferous to mid-permian, and cooler, possibly deeper marine environments in the late permian (stemmerik 2000). in east greenland, non-marine conditions prevailed during the entire carboniferous and, following a prolonged, early permian hiatus, warm-water carbonates were deposited during the late permian. stemmerik (2000) concluded that the changes reflect large-scale shifts in palaeoclimatic and subsidence patterns related to the northward drift of the area and ongoing rifting in the region. stemmerik & worsley (2005) correlated the late palaeozoic stratigraphy across north greenland, svalbard and the barents shelf, and found evidence for regional, mid-late carboniferous uplift and erosion and local, permian uplift and erosion. for a long time, it was difficult to correlate the mesozoic succession both within the wandel sea basin and with basins in surrounding areas (dypvik et al. 2002; håkansson & pedersen 2015). however, a new jurassic– cretaceous lithostratigraphy for kilen by hovikoski et al. (2018) showed correlations between kilen and northeast greenland, svalbard and the barents shelf. a triassic lithostratigraphy for the wandel sea basin presented evidence of a shallow shelf to deep shelf or basin floor transect that forms an outcrop analogue to contemporaneous intervals in the western barents sea basins and the offshore danmarkshavn basin (bjerager et al. 2019). these observations support the suggestion that the now isolated components of the wandel sea basin were once parts of a more extensive depositional system. 7 conclusions the results of this study reveal a long history of episodic burial and exhumation across the wandel sea basin. the earliest episode recognised is a phase of early permian uplift and erosion that led to the formation of the basin-wide hiatus between the upper permian and  upper carboniferous strata. the exhumation removed 2–3 km of rocks across the basin, including a cover of silurian to carboniferous sediments on herluf trolle land. late triassic uplift and erosion, which is likely to have affected the entire basin, led to formation of a hiatus between the upper and middle triassic formations on herluf trolle land when kilometre-scale thick covers of sediments were removed, including a cover of more than 2.7 km of upper carboniferous to triassic sediments on holm land. late jurassic uplift and erosion probably affected the entire basin and led to formation of a hiatus between the upper jurassic and upper triassic strata involving removal of kilometre-thick sedimentary covers, for k r o n p r i n s c h r i s t i a n l a n d ø hovgaard p e a r y l a n d 48 kctz hffz tlfz 500 100 km 2.5 km paleocene– eocene (gc1113-9, -10) >2.7 km upper cretaceous – paleocene (gc1113-5) present extent of the carboniferous–palaeogene deposits of the wandel sea basin >2.7 km carboniferous– triassic (gc1113-18) >3.3 km cretaceous– paleocene (gc1113-58) >3 km triassic – middle jurassic (gc1113-22) 40ºw 35ºw 30ºw 25ºw 20ºw 15ºw 10ºw 35ºw 30ºw 25ºw 20ºw 15ºw 80 ºn 82 ºn 81 ºn 83 ºn fig. 20 present-day outline of the carboniferous–palaeogene wandel sea basin compared with estimated thicknesses of removed covers (fig. 18). substantial thicknesses of sediments of mesozoic and palaeogene age have been present within the present outline of the basin, implying that they extended beyond the present-day outline of the wandel sea basin. the conversion of palaeotemperatures to burial depths is based on an assumed palaeogeothermal gradient of 30°c/km and palaeosurface temperatures of 10°c for the end-eocene episode and 20°c for earlier episodes. hffz: harder fjord fault zone. kctz: kap canon thrust zone. tlfz: trolle land fault zone. geology legend and symbols defined in fig. 4. https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45. 5298. https://doi.org/10.34194/geusb.v45.5298 40 of 43 www.geusbul let in.org example, more than 3 km of triassic to middle jurassic sediments on herluf trolle land. mid-cretaceous cooling is identified in only one sample within the wandel sea basin, but this episode is contemporaneous with regional exhumation in north-east greenland. we, therefore, suggest that both areas were affected by exhumation at this time, corresponding to the aptian to turonian hiatus on herluf trolle land. mid-paleocene cooling and exhumation define the timing of compression along the major fault zones (e.g. the kronprins christians lands orogeny). a cover of more than 3 km of upper cretaceous to paleocene sediments was present on the north coast of kronprins christian land prior to the paleocene exhumation. we thus agree with previous interpretations that the high thermal maturities in the southern part of kilen reflect maximum burial prior to basin inversion. we conclude that the mid-paleocene inversion of the major fault zones in the wandel sea basin took place during the first stage of the eurekan orogeny after the onset of sea-floor spreading west of greenland. afta data from most samples analysed from the northern part of kronprins christian land show evidence of paleocene cooling, including samples from around nakkehoved where upper cretaceous rocks show extreme maturities. cooling from the extreme heating must have occurred in late cretaceous or paleocene times, as afta data define only minimum estimates of the maximum paleotemperature. given the complexity of the region around nakkehoved, additional sampling for afta and vr combined with stratigraphic and sedimentological studies is required before the cause of the extreme thermal imprint there can be understood. regional exhumation of the wandel sea basin began at the end of the eocene and led to removal of most of the kilometre-thick sedimentary cover that had been deposited during eocene subsidence. peary land, north of the hffz, was uplifted about 1 km more than the area south of the fault during this episode, suggesting a tectonic offset. the regional denudation and reverse faulting that began at the end of the eocene took place after sea-floor spreading west of greenland had ended and thus post-date the eurekan orogeny. the results of this study thus define two distinct episodes of palaeogene cooling and exhumation across the wandel sea basin: (1) mid-paleocene cooling identified mainly in inverted fault zones and (2) cooling that began at the end of the eocene, which affected a wide area and reflects regional exhumation following deeper burial. mid–late miocene cooling is identified from afta data in near-coastal locations. this episode coincides with a phase of late miocene uplift and incision that initiated the formation of the present-day landscape in west and north-east greenland. miocene palaeotemperatures in the present study area are thus likely to reflect a similar development. the preserved sedimentary sequences of the wandel sea basin represent remnants of once much thicker strata removed in separate exhumation episodes. the thicknesses of these sedimentary covers imply that they must have extended substantially beyond the present-day boundary of the basin. the wandel sea basin may, thus, at times have been coherent with the carboniferous–palaeogene basins in north-east greenland, svalbard and the barents sea. acknowledgements many colleagues from the geological survey of denmark and greenland (geus) and from the university of copenhagen contributed with samples to this study, more than we can list here. niels henriksen (oscar) provided invaluable help in retrieving samples collected in the past few decades from geus’ core store. john r. hopper, kristian svennevig and morten bjerager kindly provided with comments on the manuscript. reviewers andrew carter and eckart håkansson provided constructive comments in the manuscript. additional information funding this research work was funded by a consortium of oil companies and geus. competing interests the authors declare no competing interests. author contributions pj: conceptualisation, funding acquisition, data curation, investigation, writing – review & editing. pfg: formal analysis, methodology, investigation, writing – review & editing jac: investigation, writing – review & editing. additional files two supplementary files are available at https://doi.org/10.22008/fk2/ vlo6nq references alsen, p. mcroberts, c., svennevig, k., bojesen-koefoed, j., hovikoski, j. & piasecki, s. 2017: the isrand formation: a middle triassic daonella-bearing, black shale unit in kilen, north greenland (with a note on the triassic in amdrup land). newsletters on stratigraphy 50, 31–46. https://doi.org/10.1127/nos/2016/0341 armstrong, h., smith, m., aldridge, r. & tull, s. 1994: thermal maturation of the lower palaeozoic strata of northern greenland from conodont colour alteration index (cai) data: implications for burial history and hydrocarbon exploration. geological magazine 131, 219–230. https:// doi.org/10.1017/s0016756800010748 arne, d., grist, 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(eds): strike-slip deformation, basin formation, and sedimentation. the society of economic paleontologists and mineralogists (sepm) special publication 37, 339–359. https://doi.org/10.2110/pec.85.37.0339 stemmerik, l. 2000: late palaeozoic evolution of the north atlantic margin of pangea. palaeogeography, palaeoclimatology, palaeoecology 161, 95–126. https://doi.org/10.1016/s0031-0182(00)00119-x stemmerik, l. & worsley, d. 2005: 30 years on — arctic upper palaeozoic stratigraphy, depositional evolution and hydrocarbon prospectivity. norwegian journal of geology 85, 151–168. stemmerik, l., dalhoff, f., larsen, b.d., lyck, j.m., mathiesen, a. & nilsson, i. 1998: wandel sea basin, eastern north greenland. geology of greenland survey bulletin 180, 55–62. https://doi.org/10.34194/ggub.v180.5086 surlyk, f. 1990: timing, style and sedimentary evolution of late palaeozoic – mesozoic extensional basins in east greenland. in: hardman, r.f.p. & brooks, j. (eds) tectonic events responsible for britain’s oil and gas reserves. geological society special publications (london) 55, 107–125. https://doi.org/10.1144/gsl.sp.1990.055.01.05 surlyk, f. 1991: tectonostratigraphy of north greenland. bulletin grønlands geologiske undersøgelse 160, 27–47. svennevig, k., guarnieri, p. & stemmerik, l. 2016: tectonic inversion in the wandel sea basin: a new structural model of kilen (eastern north greenland). tectonics 35, 2896–2917. https://doi.org/10.1002/2016tc004152 svennevig, k., guarnieri, p. & stemmerik, l. 2017: 3d restoration of a cretaceous rift basin in kilen, eastern north greenland. norwegian journal of geology 97, 21–32. https://doi.org/10.17850/ njg97-1-02 svennevig, k., alsen, p., guarnieri, p., hovikoski, j., wesenberg lauridsen, b., krarup pedersen, g., nøhr-hansen, h. & sheldon, e. 2018: descriptive text to the geological map of greenland, 1:100 000, kilen 81 ø.1 syd. geological survey of denmark and greenland map series 8, 1–29. https://doi.org/10.34194/geusm.v8.4526 tegner, c., storey, m., holm, p.m., thorarinsson, s.b., zhao, x., lo, c.h. & knudsen, m.f. 2011: magmatism and eurekan deformation in the high arctic large igneous province: 40ar-39ar age of kap washington group volcanics, north greenland. earth and planetary science letters 303, 203–214. https://doi.org/10.1016/j.epsl.2010.12.047 tessensohn, f. & piepjohn, k. 2000: eocene compressive deformation in arctic canada, north greenland and svalbard and its plate tectonic causes. polarforschung 68, 121–124. vamvaka, a., pross, j., monien, p., piepjohn, k., estrada, s., lisker, f. & spiegel, c. 2019: exhuming the top end of north america: episodic evolution of the eurekan belt and its potential relationships to north atlantic plate tectonics and arctic climate change. tectonics 38, 4207–4228. https://doi.org/10.1029/2019tc005621 von gosen, w. & piepjohn, k. 2003: eurekan transpressive deformation in the wandel hav mobile belt (northeast greenland). tectonics 22, 28. https://doi.org/10.1029/2001tc901040 https://doi.org/10.34194/geusb.v45.5298 http://www.geusbulletin.org https://doi.org/10.2110/pec.85.37.0339 https://doi.org/10.1016/s0031-0182(00)00119-x https://doi.org/10.34194/ggub.v180.5086 https://doi.org/10.1144/gsl.sp.1990.055.01.05 https://doi.org/10.1002/2016tc004152 https://doi.org/10.17850/njg97-1-02 https://doi.org/10.17850/njg97-1-02 https://doi.org/10.34194/geusm.v8.4526 https://doi.org/10.1016/j.epsl.2010.12.047 https://doi.org/10.1029/2019tc005621 https://doi.org/10.1029/2001tc901040 thermo-tectonic development of the wandel sea basin, north greenland abstract 1 introduction 2 geological setting 2.1 stratigraphic record 2.1.1 palaeozoic strata 2.1.2 mesozoic strata 2.1.3 palaeogene strata 2.2 structural elements of the wandel sea basin and their relation to the eurekan orogeny 3 apatite fission-track analysis 3.1 new afta data from northern and eastern greenland 3.2 thermal history interpretation based on afta data 3.2.1 principles of thermal history interpretation 3.2.2 thermal history interpretation: results 3.3 definition of major regional palaeothermal episodes 3.4 regional variation in palaeothermal episodes 4 thermal history insights from maturity data 4.1 maturity of lower palaeozoic strata 4.2 maturity of mesozoic-palaeogene strata 4.3 comparison of afta and vr data 4.3.1 northern peary land 4.3.2 herluf trolle land 4.3.3 east of herluf trolle land 4.4 synthesis 5 geological controls on the cooling episodes identified from afta 5.1.palaeozoic-mesozoic cooling events 5.1.1 early permian episode 5.1.2 late triassic episode 5.1.3 late jurassic episode 5.1.4 mid-cretaceous episode 5.2 cenozoic cooling episodes 5.2.1 mid-paleocene episode 5.2.2 end-eocene episode 5.2.3 mid-late miocene episode 6 discussion 6.1 mid-paleocene onset of the eurekan orogeny 6.2 recent results from arctic canada 6.3 end-eocene tectonics in relation to the eurekan orogeny 6.4 the extreme heating recorded at nakkehoved 6.5 previous extent of the wandel sea basin 7 conclusions acknowledgements additional information references figures fig. 1 cenozoic sea-floor spreading in the northern north atlantic area. sea-floor spreading between north america and greenland (62–55 ma) caused the first stage of the eurekan orogeny. during this time, greenland moved together with europe. spreading between europe and greenland (55–35 ma) resulted in the second stage of the eurekan orogeny. during this time, north america, greenland and europe moved as separate plates. sea-floor spreading ceased west of greenland (at c. 35 ma) after which greenland moved together with north america and the eurekan orogeny ended. the danmarkshavn basin is located off north-east greenland between 76 and 82°n. lr: lomonosov ridge. wsb: wandel sea basin. modified from brozena et al. (2003), oakey & chalmers (2012) and gaina et al. (2017). fig. 2 outline of the high arctic adjacent to north greenland. extent of kronprins christian land orogeny according to håkansson & pedersen (2001) and this study. lr: lomonosov ridge. detail of fig. 1 with additions from von gosen & piepjohn (2003), oakey & chalmers (2012) and piepjohn et al. (2016). fig. 3 elevation and place names of the study area in eastern north greenland. pr: prinsesse. fig. 4 geology of the carboniferous–palaeogene wandel sea basin and surrounding areas. the remnants of the carboniferous to palaeogene sediments occur between northern peary land and kronprins christian land and thus define the extent of the basin. kap washington group is of campanian to earliest palaeogene age (håkansson & pedersen 2015). håkansson & pedersen (2015) indicated the trolle land fault zone as one continuous zone from herluf trolle land to kronprins christian land. here, we indicate one possible correlation between these two areas. based on escher & pulvertaft 1995 with modifications after croxton et al. 1980; hovikoski et al. 2018; piasecki et al. 2018. egfz: east greenland fracture zone. hffz: harder fjord fault zone. kctz: kap canon thrust zone. kr: kap rigsdagen. pr.m.ø: prinsesse margrethe ø. pr.t.ø: prinsesse thyra ø. tlfz: trolle land fault zone. fig. 5 wandel sea basin stratigraphy and formation names. the representation of the mixed volcanic, volcano-clastic and lacustrine deposits of the kap washington group has been simplified. some formation names are not shown; the upper permian kap kraka fm in peary land (håkansson & pedersen 2015) and the pre-late bajocian (?) gletscherport fm in the kilen area (hovikoski et al. 2018). compiled after bjerager et al. 2019; gautier et al. 2011; håkansson & pedersen 2015; hopper et al. 2014; hovikoski et al. 2018; ineson et al. 2020; piasecki et al. 2018; stemmerik et al. 1998; svennevig et al. 2018; tegner et al. 2011. bf: birkelund fjeld. dd: dromledome. k.c. land: kronprins christian land. kl: kuglelejet. kw: kap washington group. pb: parish bjerg. penns: pennsylvanian. fig. 6 location of samples that yielded apatite and division of the study area into six geographical areas with similar thermal histories used to define the timing of key episodes in fig. 10. a: areas 1–5. b: area 6. geology legend in fig. 4. egfz: east greenland fracture zone. hffz: harder fjord fault zone. kctz: kap canon thrust zone. kr: kap rigsdagen. pr.t.ø: prinsesse thyra ø. tlfz: trolle land fault zone. fig. 7 apatite fission-track ages in outcrop samples from eastern north greenland (data in table 2). youngest ages <60 ma are focussed along the north coast, while older ages are found in inland samples. egfz: east greenland fracture zone. hffz: harder fjord fault zone. kctz: kap canon thrust zone. kr: kap rigsdagen. pr.t.ø: prinsesse thyra ø. tlfz: trolle land fault zone. geology legend in fig. 4. fig. 8 relationship between the mean track length and fission-track age. a: outcrop samples from eastern north greenland (this study). error bars are shown as ± 2 sigma. b: results from this study compared with those from north-east greenland (japsen et al. in press). ages span a similar range in both regions. however, data from this study are characterised by higher mean track lengths than those from north-east greenland, which probably reflects the lower degree of miocene cooling in samples from this study. error bars are omitted for clarity. fig. 9 illustration of thermal history solutions extracted from afta data. results are presented in terms of up to three palaeothermal episodes, that is, times when a rock sample was hotter than it is today. note that the palaeotemperature constraint t1 is a minimum estimate because this is the event in which the apatites begin to retain tracks. in most situations, three events are the most that can be defined from afta due to various factors, including the natural spread in track lengths of a single population of tracks and the rapid decrease in the rate of annealing across the range from c. 110°c to below 60°c. t1, t2 and t3: time intervals during which cooling from the peak palaeotemperature began. t1, t2 and t3: palaeotemperature intervals. c.i.: confidence interval. colours are used to illustrate attribution to regional episodes, although those used here are purely schematic. fig. 10 timing constraints on the onset of cooling for discrete cooling episodes derived from the afta in each outcrop sample. thermal history solutions for each sample are listed in table 3. horizontal-shaded bars represent 95% confidence intervals on the time of cooling onset, as illustrated in fig. 9. samples are divided into six areas (areas 1–6; fig. 6). assuming these results represent regionally synchronous cooling episodes, results across the entire region require seven separate episodes, as shown by vertical shaded columns. constraints for the onset of each episode are listed at the bottom of the figure (see table 4). timing constraints for individual samples are assigned to a common regional event when adjacent samples cool from similar palaeotemperatures with consistent timings. in some cases, the estimated onset of cooling pre-dates the depositional age of the sample, indicating cooling of the apatites in sediment provenance regions prior to deposition in the host sample. fig. 11 palaeotemperatures derived from afta. a: early permian (295–290 ma). b: late triassic (225–210 ma). c: late jurassic (165–150 ma). d: paleocene (c. 60 ma). e: end-eocene (c. 35 ma). f: middle to late miocene (15–5 ma). palaeotemperature for the mid-cretaceous episode (115–90 ma) is not shown as only one sample falls within the map frame. the small arrow in the lower right corner of a, e and f indicates sample gc1113-48, which is located just south of the figure frame. geology legend in fig. 4. numbers next to afta samples denote sample numbers. hffz: harder fjord fault zone. kctz: kap canon thrust zone. kr: kap rigsdagen. pr.t.ø: prinsesse thyra ø. tlfz: trolle land fault zone. fig. 12 published vitrinite reflectance (vr) values from mesozoic and palaeogene units of the wandel sea basin. coloured boxes: vr values and corresponding maximum palaeotemperatures (burnham & sweeny 1989). source of vr data indicated by letters a: paech & estrada (2018), b: pedersen et al. (2018), c: håkansson et al. (1994). when values are available from multiple sources, those from paech & estrada (2018) are shown. pr.d.ø: prinsesse dagmar ø. pr.m.ø: prinsesse margrethe ø. pr.t.ø: prinsesse thyra ø. detail of kilen in fig. 15. fig. 13 cenozoic palaeothermal constraints for herluf trolle land. heating during the paleocene event (c. 60 ma) is controlled by the numerous faults belonging to the trolle land fault zone, as suggested by paech & estrada (2018). only samples north-east of ‘fault f’ were affected by the paleocene thermal event. white boxes: thermal history solutions from afta data; colour coding as in fig. 10. coloured boxes: vr values and corresponding maximum palaeotemperatures from adjacent samples; see fig. 12. source of vr data indicated by letter a: paech & estrada (2018). map location indicated in fig. 12. basemap modified from bjerager et al. (2019). fig. 14 cenozoic palaeothermal constraints east of herluf trolle land. white boxes: thermal history solutions from afta, colour coding as in fig. 10. coloured boxes: vr values and corresponding maximum palaeotemperatures from adjacent samples, colour coding as in fig. 12. source of vr data indicated by letters a: paech & estrada (2018), c: håkansson et al. (1994). solid black lines: faults. map location indicated in fig. 12. detail of kilen in fig. 15. key to bedrock geology and other abbreviations in fig. 4. fig. 15 palaeothermal constraints for south-east kilen. the palaeotemperature constraints from vr are consistent with those from afta, but the timing constraints on the onset of cooling vary. white boxes: thermal history solutions from afta, colour coding as in fig. 10. numbers next to afta samples denote sample numbers. blue boxes: vr values and associated maximum palaeotemperatures for cretaceous sediments. source of vr data indicated by letters b: pedersen et al. (2018) and c: håkansson et al. (1994). map location in fig. 14. base map modified from hovikoski et al. (2018, fig. 4) where further details of the map are explained. fig. 16 comparison between the regional cooling episodes identified from afta (table 4) and the stratigraphic scheme for the wandel sea basin (fig. 5). the thickness of the horizontal, coloured bars indicates the interval during which the cooling began in each episode. fig. 17 timing of regional, post-devonian episodes of uplift and erosion in three regions of greenland estimated from afta. a: central west greenland (japsen et al. 2006, 2009). b: eastern north greenland (table 4; this study). c: north-east greenland (japsen et al. in press). break-up west of greenland took place in mid-paleocene (c. 62 ma), but sea-floor spreading ceased there by the end of the eocene. in contrast, the opening of the north-east atlantic, east of greenland, began at the paleocene–eocene transition (c. 56 ma) and is still ongoing (chalmers & pulvertaft 2001; oakey & chalmers 2012; gaina et al. 2017). the fram strait between greenland and svalbard opened in early miocene (c. 20 ma; jokat et al. 2016). l: late. e: early. m: middle. c: carboniferous. pe: permian. tr: triassic. j: jurassic. kl: lower cretaceous. ku: upper cretaceous. pal: paleocene. eo: eocene. oli: oligocene. mio: miocene. pl: pliocene. q: quaternary. fig. 18 thermal history diagrams illustrating the heating and cooling (burial and exhumation) of four samples that provide stratigraphic constraints on the age of removed sections. a: gc1113-18 (20 m a.s.l.) of upper carboniferous kap jungersen fm exposed on holm land. the sample reached a palaeotemperature above 100°c prior to late triassic cooling. this palaeotemperature corresponds to burial below a cover of carboniferous to triassic sediments, with a thickness exceeding 2.7 km. b: gc1113-22 (200 m a.s.l.) of lower triassic parish bjerg fm exposed on herluf trolle land. the sample reached a palaeotemperature above 115°c prior to late jurassic cooling, corresponding to burial below a cover of triassic to middle jurassic sediments with a thickness exceeding 3.1 km.the conversion of palaeotecmperatures to burial depths is based on an assumed palaeogeothermal gradient of 30°c/km and palaeosurface temperatures of 10°c for the end-eocene episode and 20°c for earlier episodes. red triangles: constraints for conditions at the surface during deposition of sediments or at the present day. coloured boxes: constraints from afta on palaeotemperature prior to cooling (the width of the boxes is fixed for better overview). boxes extend below the x-axis of the diagram where only a lower limit of the palaeotemperature is known (e.g. >110°c). sample locations in fig. 6. fig. 19 3d-elevation model of the region around independence fjord (globe 1999; map location in fig. 3). note that independence fjord and wandel dal (valley in danish) are incised below a low-relief surface at about 1 km a.s.l. this is in contrast to more elevated and dissected, alpine relief north-west of independence fjord. the width of the diagram is about 250 km. illustration: johan m. bonow. fig. 20 present-day outline of the carboniferous–palaeogene wandel sea basin compared with estimated thicknesses of removed covers (fig. 18). substantial thicknesses of sediments of mesozoic and palaeogene age have been present within the present outline of the basin, implying that they extended beyond the present-day outline of the wandel sea basin. the conversion of palaeotemperatures to burial depths is based on an assumed palaeogeothermal gradient of 30°c/km and palaeosurface temperatures of 10°c for the end-eocene episode and 20°c for earlier episodes. hffz: harder fjord fault zone. kctz: kap canon thrust zone. tlfz: trolle land fault zone. geology legend and symbols defined in fig. 4. tables table 1 sample details. a digital version of this table is available in online supplementary file s1 table 2 afta data. a digital version of this table is available in online supplementary file s1 table 3 thermal history solutions. a digital version available in online supplementary file s1 table 4 intervals defining the onset of episodes of cooling based on afta data in all samples and the possible origin of the cooling during each episode untitled 1 geological survey of denmark and greenland bulletin 8 · 2005 structural analysis of the rubjerg knude glaciotectonic complex, vendsyssel, northern denmark stig a. schack pedersen geological survey of denmark and greenland danish ministry of the environment 2 geological survey of denmark and greenland bulletin 8 keywords northern jylland, denmark, weichselian, glacial geology, glaciotectonics, thin-skinned thrust faulting, balanced cross-section, thrustfault dynamics, imbricate duplexes, mud diapirs, piggyback basins. cover the coastal clif f (99 m high at its highest point) at rubjerg knude on the west coast of vendsyssel, northern denmark. the lower two-thirds of the cliff, beneath the prominent dark sub-horizontal surface, forms part of the cross-section through the rubjerg knude glaciotectonic complex displaying imbricated thrust sheets composed of the lønstrup klint formation (bluish-grey colour) and the overlying rubjerg knude formation (yellow colour), both of late weichselian age. the thrust sheets are truncated by a glaciotectonic unconformity (the prominent surface), upon which the kattegat till formation is only preserved as a boulder bed due to subsequent aeolian erosion of the till matrix. the upper third of the clif f comprises recent aeolian dune sands that have accreted over the last 100 years and now encroach on the rubjerg knude lighthouse, the top of which is just visible above the clif ftop. photo: stig a. schack pedersen (august 1984). chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, geological institute, university of aarhus; minik rosing, geological museum, university of copenhagen; finn surlyk, geological institute, university of copenhagen scientific editor of this volume: jon r. ineson editorial secretaries: esben w. glendal and birgit eriksen illustrations: benny m. schark and alice rosenstand digital photographic work: benny m. schark graphic production: knud gr@phic consult, odense, denmark printers: schultz grafisk, albertslund, denmark manuscript submitted: 8 august 2003 final version approved: 11 february 2005 printed: 15 december 2005 this monograph has been accepted by the faculty of natural sciences, university of copenhagen, for public defence of the degree of doctor of science. issn 1604-8156 isbn 87-7871-168-1 geological survey of denmark and greenland bulletin the series geological survey of denmark and greenland bulletin replaces geology of denmark survey bulletin and geology of greenland survey bulletin. citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 8, 192 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk or geografforlaget aps rugårdsvej 55, dk-5000 odense c, denmark phone: +45 63 44 16 83, fax: +45 63 44 16 97, e-mail: go@geografforlaget.dk © danmarks og grønlands geologiske undersøgelse (geus), 2005 3 contents abstract . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9 introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 11 history of the present investigation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 11 objectives . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 14 glacial tectonics – concepts and models . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 14 previous conceptual models . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 14 thin-skinned thrust faulting: the concept . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16 thrust-fault modelling . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16 test model 1. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19 test model 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19 test model 3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 21 test model 4 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 21 test models: concluding remarks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 22 concept of balanced cross-section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 22 location and construction of the rubjerg knude cross-section . . . . . . . . . . . . . . . . . 23 location. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23 photogrammetric work. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23 digital editing . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25 construction of the balanced cross-section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 26 geological setting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 27 lithostratigraphy. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 33 skærumhede group . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 33 stortorn formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 35 lønstrup klint formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 39 upper weichselian lithostratigraphic units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 43 rubjerg knude formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 43 kattegat till formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 48 ribjerg formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 50 mid danish till formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 53 vendsyssel formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 55 structural description of sections . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 60 ulstrup section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 60 tectonic architecture . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 61 sedimentary units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 62 lønstrup klint formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 62 rubjerg knude formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 63 structures and breccias . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 63 thrust-zone breccias . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 64 foreland-dipping hanging-wall flat faults . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 64 collapse structure in the ulstrup rende . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 66 interpretation of structural development . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 66 stensnæs section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 70 tectonic architecture . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 70 sedimentary units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 72 lønstrup klint formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 72 rubjerg knude formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 73 structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 73 4 imbricate duplex folding . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 73 extensional faults . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 73 interpretation of structural development . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 78 martørv bakker section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 78 tectonic architecture . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 79 sedimentary units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 80 lønstrup klint formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 81 rubjerg knude formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 81 structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 82 imbricate duplexes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 82 normal fault . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 82 hydrodynamic brecciation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 82 interpretation of structural development . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 82 kramrende section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 84 tectonic architecture . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 85 sedimentary units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 86 lønstrup klint formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 87 rubjerg knude formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 87 structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 87 thrust faults . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 87 kramrende diapir . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 88 reverse faults . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 88 interpretation of structural development . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 89 brede rende section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 90 tectonic architecture . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 90 sedimentary units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 93 lønstrup klint formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 93 rubjerg knude formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 94 structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 94 diapir structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 96 brede rende normal fault . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 97 frost wedges . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 97 interpretation of structural development . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 97 sandrende section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 99 tectonic architecture . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 99 sedimentary units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 101 lønstrup klint formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 101 rubjerg knude formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 101 structures and breccias . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 103 normal faults . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 103 diapir structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 104 frost wedge . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 104 interpretation of structural development . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 104 stenstue rende section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 105 tectonic architecture . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 106 sedimentary units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 107 lønstrup klint formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 107 rubjerg knude formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 108 structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 108 thrust-fault structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 108 hanging-wall anticlines . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 109 5 normal faults . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 110 slump folding . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 110 interpretation of structural development . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 110 grønne rende section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 111 tectonic architecture . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 112 sedimentary units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 112 lønstrup klint formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 112 rubjerg knude formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 112 structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 113 gr01 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 113 gr02 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 114 gr03 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 114 gr04 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 114 gr05 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 115 gr06 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 115 gr07 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 116 gr08 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 116 gr09 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 116 gr10 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 116 gr11 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 117 gr12 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 117 gr13 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 117 interpretation of structural development . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 117 rubjerg knude fyr section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 118 tectonic architecture . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 118 sedimentary units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 121 structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 122 anastomosing thrust-fault brecciation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 122 interpretation of structural development . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 122 stortorn section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 122 tectonic architecture . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 123 st01 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 123 st02 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 123 st03 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 123 st04 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 123 st05 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 124 st06 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 124 st07 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 124 st08 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 124 st09 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 125 st10 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 125 sedimentary units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 126 structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 126 interpretation of structural development . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 126 moserende section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 128 tectonic architecture . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 128 mr01 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 129 mr02 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 130 mr03 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 130 mr04 and mr05 thrust sheets . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 130 mr06–mr08 thrust sheets . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 131 6 mr09 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 132 mr10 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 132 mr11 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 132 mr12 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 132 mr13 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 133 sedimentary units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 134 lønstrup klint formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 134 rubjerg knude formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 134 structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 134 diapir structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 135 thrust faults . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 135 footwall synclines . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 135 interpretation of structural development . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 136 mårup kirke section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 137 tectonic architecture . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 138 mk01 thrust sheet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 139 mk02–mk04 thrust sheets . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 139 mk05–mk07 thrust sheets . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 139 mk08–mk10 thrust sheets . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 140 mk11–mk20 thrust sheets . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 140 sedimentary units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 141 structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 141 interpretation of structural development . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 141 fault-bend-fold model for duplex units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 141 characterisation of thrust duplex mk11–mk20. . . . . . . . . . . . . . . . . . . . . . . . . . . . 143 discussion of structural development . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 143 ribjerg section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 144 ‘store blå’ and ‘lille blå’. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 144 tectonic architecture . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 145 sedimentary units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 146 skærumhede group . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 146 blå-unconformity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 146 ribjerg formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 146 mid danish till formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 147 vendsyssel formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 147 structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 147 interpretation of glacial geology and stratigraphic development . . . . . . . . . . . . . . . . 147 dynamic development of the thin-skinned thrust faulting . . . . . . . . . . . . . . . . . . . . . . 148 moserende section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 148 moserende section: summary data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 149 stortorn section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 149 stortorn section: summary data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 156 rubjerg knude fyr section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 156 rubjerg knude fyr section: summary data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 156 grønne rende section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 156 grønne rende section: summary data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 157 stenstue rende section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 157 stenstue rende section: summary data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 161 sandrende section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 161 sandrende section: summary data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 163 brede rende section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 163 7 brede rende section: summary data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 166 kramrende section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 166 kramrende section: summary data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 168 martørv bakker section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 168 martørv bakker section: summary data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 170 stensnæs section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 170 stensnæs section: summary data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 172 ulstrup section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 173 ulstrup section: summary data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 174 summary of dynamic development . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 174 discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 178 thrust-fault architecture . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 178 balanced cross-section . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 179 thrust brecciation and diapirism . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 180 thrust-fault dynamics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 181 syntectonic deposition . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 181 proglacial and subglacial deformation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 182 glacial geological conditions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 183 conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 185 acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 185 references . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 186 appendix 1 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 190 thrust-fault terminology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 190 appendix 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 192 specification of photogrammetric work . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 192 8 9 abstract pedersen, s.a.s. 2005: structural analysis of the rubjerg knude glaciotectonic complex, vendsyssel, northern denmark. geological survey of denmark and greenland bulletin 8, 192 pp. the rubjerg knude glaciotectonic complex is a thin-skinned thrust-fault complex that was formed during the advance of the scandinavian ice sheet (30 000 – 26 000 b.p.); it is well exposed in a 6 km long coastal profile bordering the north sea in northern denmark. the glaciotectonic thrust-fault deformation revealed by this cliff section has been subjected to detailed structural analysis based on photogrammetric measurement and construction of a balanced cross-section. thirteen sections are differentiated, characterising the distal to proximal structural development of the complex. the deformation affected three stratigraphic units: the middle weichselian arctic marine stortorn formation, the mainly glaciolacustrine lønstrup klint formation and the dominantly fluvial rubjerg knude formation; these three formations are formally defined herein, together with the skærumhede group which includes the stortorn and lønstrup klint formations. the rubjerg knude formation was deposited on a regional unconformity that caps the lønstrup klint formation and separates pre-tectonic deposits below from syntectonic deposits above. in the distal part of the complex, the thrust-fault architecture is characterised by thin flatlying thrust sheets displaced over the footwall flat of the foreland for a distance of more than 500 m. towards the proximal part of the complex, the dip of the thrust faults increases, and over long stretches they are over-steepened to an upright position. the lowest décollement zone is about 40 m below sea level in the proximal part of the system, and shows a systematic step-wise change to higher levels in a distal (southwards) direction. the structural elements are ramps and flats related to hanging-wall and footwall positions. above upper ramp-hinges, hanging-wall anticlines developed; footwall synclines are typically related to growth-fault sedimentation in syntectonic piggyback basins, represented by the rubjerg knude formation. blocks and slump-sheets constituting parts of the lønstrup klint formation were derived from the tips of up-thrusted thrust sheets and slumped into the basins. mud diapirs are a prominent element in the thrust-fault complex, resulting from mud mobilisation mainly at hanging-wall flats and ramps. shortening during thrust-fault deformation has been calculated as 50%. only about 11% of the initial stratigraphic units subjected to thrust faulting has been lost due to erosion. the thrust-fault deformation was caused by gravity spreading of an advancing ice sheet. overpressured mud-fluid played an important role in stress transmission. the average velocity of thrust-fault displacement is estimated at 2 m per year, which led to compression of a 12 km stretch of flat-lying sediments, c. 40 m in thickness, into a thrust-fault complex 6 km in length. the thrust-fault complex is truncated by a glaciotectonic unconformity, formed when the advancing ice sheet finally overrode the complex. when this ice sheet melted away, a hilland-hole pair was formed, and meltwater deposits derived from a new ice-advance (ne-ice) filled the depression. the ne-ice overran the complex during its advance to the main stationary line situated in the north sea. when this ice in turn melted away (c. 19 000 – 15 000 b.p.), the glacial landscape was draped by arctic marine deposits of the vendsyssel formation (new formation defined herein). _________________________________________________________________________________________ author’s address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: sasp@geus.dk 10 55n 5e 10e 15n 60n scandinavian ice sheet norway denmark copenhagen göteborg møns klint bovbjerg fur knudeklint hanklint mols hoved rügen ristinge klint lønstrup klint sweden germany baltic ice advance swedish ice advance norwegian ice advance 28 ka bp 30 ka bp 20 ka bp 17 ka bp 17 ka bp 17 ka bp 28 ka bp fig. 1. map of the danish basin indicating the distribution of the scandinavian ice sheet during the three main ice advance events, with source areas in southern norway, central sweden and the baltic, in the middle–late weichselian. the approximate timing of the stationary lines are given; the early progressive ice advance is indicated in black, the subsequent late ice border lines in red. the locations of major glaciotectonic complexes formed during the ice advances are indicated by asterisks. 11 introduction glaciotectonic studies in denmark have a long tradition, and an important part of structural geology studies in denmark concern glacial tectonic deformation resulting from the southward advance of the scandinavian ice sheet in the pleistocene (fig. 1). the description of the geological structures dates back to puggaard (1851), who made one of the first extensive and detailed danish structural analyses of a tectonic complex and provided a classic cross-section of møns klint. johnstrup (1874) established the concept of glacial deformation. the next milestone in glacial tectonic studies in denmark was by jessen (1918, 1931), whose detailed survey of lønstrup klint (fig. 1) included a structural analysis and an attempt at a glaciodynamic interpretation of the deformation structures observed. the lønstrup klint coastal section includes the rubjerg knude glaciotectonic complex, which is the subject of this study (fig. 2). a danish school of glaciotectonic studies subsequently developed (madsen 1916; jessen 1931; gry 1940, 1941; rosenkrantz 1944; berthelsen 1973, 1975, 1978, 1979; sjørring 1974, 1977, 1981, 1983; rasmussen 1975; petersen 1978; houmark-nielsen 1987, 1988; pedersen 1987, 1993, 1996, 2000; pedersen & petersen 1988, 1995, 1997; pedersen et al.1988; klint & pedersen 1995; jakobsen 1996), which has naturally been stimulated by geologists working with glaciotectonic structural geology internationally (banham 1977, 1988; stephan 1980; aber 1982, 1993; ehlers 1983; van der wateren 1985, 1992; boulton 1986; boulton & hindmarsh 1987; croot 1987, 1988; meer 1987; goldthwait & matsch 1988; aber et al. 1989; hart 1990; hart & watts 1997; bennett 2001). the similarity in structural geometry between glaciotectonic terrains and orogenic belts has led to prolonged debate. are glaciotectonic terrains scale models for orogenic deformation? or does the soft and synsedimentary nature of glaciotectonics differ in principle from that of fold belt deformation? arguments for deformational similarity have been put forward by berthelsen (1978, 1979), banham (1988), aber et al. (1989), van der wateren (1992) and pedersen (1987, 2000). these structural geologists share the opinion that the terminology of structural geology related to orogenic belts is applicable in the description and discussion of glaciotectonic complexes. the main differences between deformation in metamorphically altered rocks and glaciotectonic deformation of soft sediments are: (1) the presence of ‘free’ water, which enables liquefaction and fluidisation, (2) the velocity of the deformation, and (3) the shallowness of penetrative deformation. in contrast, deformation of metamorphic rocks commonly involves alteration and recrystallisation of minerals, processes that never apply to glaciotectonics. the advantage of a study of glaciotectonic complexes is that the structures are at a scale that allows them to be studied in a single exposure, in contrast to fold belts where extensive field mapping and expensive geophysical investigations are typically required for adequate documentation of the structures. furthermore, many glaciotectonic complexes are geologically young, which means that the upper structural levels are still preserved and interpretation of the full dynamic development of structural complexes is possible. the structural architecture of glaciotectonic complexes may therefore serve as inspiration for the interpretation of thin-skinned structural relationships in fold belts and thrust-fault deformation terrains. the structural analysis of the rubjerg knude glaciotectonic complex is presented as a mesoscopic model of a thinskinned thrust-fault complex (plates 1, 2). history of the present investigation this study focuses on the structural framework and dynamic development of the glacial tectonic thrustfault complex at rubjerg knude, lønstrup klint. it is based on twenty years of investigations of the lønstrup klint cliff section. the author took up the study of glacial tectonic thrust-fault structures after having concluded a ph.d. thesis on thin-skinned thrust faulting in the north greenland fold belt (pedersen 1979, 1981, 1982, 1986a, 1987). a large part of the study of the fold belt structures in peary land, north greenland, was photogrammetric mapping (pedersen 1979, 1981), undertaken at a time when geological mapping by computer-assisted photogrammetry was under development in copenhagen. this project was an integrated collaboration between the geological survey of greenland, the institute of surveying and photogrammetry of the technical university of denmark (dtu), the geological museum (gm) and the geological institute (gi) of the university of copenhagen. in the years 12 l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l diamictite glaciolacustrine and glaciofluvial sand and gravel mobilised mud non-marine sand non-marine clay and silt marine clay rubjerg knude formation lønstrup klint formation stortorn formation 0 500 1000 1500 2000 2500 3000 0 100 200 300 400 5 150014001300120011001000 2000 2100 2200 2300 2400 2500 350034003300320031003000 4100 4200 4300 4400 4500 560055005400530052005100 steps ribjerg moserende kramrende stenstue rende stortorn grønne rende brede rende stensnæs fig. 2. geological cross-section of the rubjerg knude glaciotectonic complex. for details and legend, see plate 1. 13 l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l l aeolian dunes holocene peat marine clay and sand sandy till glaciofluvial sand { mid danish till formation & kattegat till formation vendsyssel formation ribjerg formation top of dunes clifftop – glacial abrasion surface thrust, fault unconformity intraformational bedding 3500 4000 4500 5000 5500 6000 m 500 600 700 800 900 1000 m 2000 m1900180017001600 2600 2700 2800 2900 3000 m 4100 m40003900380037003600 4600 4700 4800 4900 5000 5100 m 6000 m590058005700 steps rubjerg knude fyr sandrende martørv bakkeroddervej ulstrup ulstrup rende tvonnet rende mårup kirke 14 up to 1990, techniques of geological mapping and construction of geological cross-sections based on multimodel photogrammetric analysis were developed and made available at dtu (dueholm 1992). initial investigations in co-operation with k. dueholm (dtu) and a.k. pedersen (gm) proved the applicability of multimodel photogrammetry in the study of glaciotectonic cross-sections in denmark by an examination of the møns klint clif f section (pedersen 2000). subsequently, the photogrammetric investigation of the rubjerg knude clif f section was initiated, and forms the basis of the present work. objectives the objectives of the study of the rubjerg knude glaciotectonic complex can be summarised as follows. 1. a description of an exceptionally well-exposed glaciotectonic complex, which can be taken as an example of a very low friction thrust-fault wedge, presented as a detailed cross-section based on multimodel photogrammetric measurements of the rubjerg knude clif f section. 2. a demonstration of the techniques of balanced cross-section construction that permit interpretation of the unexposed parts of the thrust-fault complex. 3. the construction of a model for the dynamic development of the proglacial thrust system that demonstrates the sequential evolution of increasing deformation intensity and the interplay with syntectonic depositional processes. 4. an interpretation of deformation processes within the framework of danish glacial stratigraphy in the late pleistocene (late middle to late weichselian c. 30 000 – 20 000 years b.p.). previous conceptual models the basic concept of glacial processes acting as the deformation agent was formulated by johnstrup (1874). his concept was primarily focused on the formation of the spectacular cliffs at møns klint in south-eastern denmark and on rügen in north-eastern germany. however, subsequently johnstrup (1882) also included the formation of the steeply inclined floes exposed in the lønstrup klint clif f section in the classic examples of glacial deformation in denmark. (the term floes is frequently used in the old glacial geology literature inspired by the idea that the dislocated sheets were groundor permafrozen; in a structural geological context, floes are identical to thrust sheets or thrustsheet segments.) johnstrup’s main conclusions concerning the glaciotectonic origin of the deformation at lønstrup klint were: (1) the dislocations are superficial without extending down to a deep root zone, and are restricted to surface phenomena, (2) the direction of movement indicated from the dip of the dislocated floes corresponds to a uniform direction of ice advance, and (3) the dislocated floes formerly constituted one undisturbed area. the detailed mapping and construction of the cross-section was presented by jessen (1918) in his geological description of the vendsyssel map sheet. however, the final detailed description of the dislocations at lønstrup klint was published later (jessen 1931). in 1927, george slater included a study of the lønstrup klint section as part of his thesis for a d.sc. degree at the university of london, which also included a study of glacial deformation at møns klint. the most striking conclusion was that the glacial deformation at lønstrup klint was caused by englacial deformation. slater (1927, p. 312) summarised thus: “… 2. the deposits represent the final positions of englacial material after the melting of the interstitial ice. 3. the type of structure is analogous to that seen in decaying arctic glaciers, and is due to the arresting of movement of the frontal part of an overloaded ice-sheet. 4. the structure has been built up in the reverse direction to the line of movement.” slater (1927) interpreted the lønstrup klint section as a variety of glacial tectonics he termed ‘the stagnant-glacier type’. subsequently, axel jessen and karl gripp exchanged ideas about proglacially formed glaciotectonic structures, and concluded that the structures jessen had observed at lønstrup klint were similar to those that gripp (1929) described from the foreland of the adglacial tectonics – concepts and models 15 vancing holmströms gletscher on spitsbergen. in his detailed and comprehensive description of his investigations, jessen (1931) concluded that the disloctions cannot have formed englacially, but must be the result of pressure building up due to loading at the margin of the advancing ice. this pressure spreads out laterally into the clayey units, which in the foreland react by splitting up into fractured dislocation sheets compressed in front of the advancing ice masses. jessen (1931) also discussed the dif ficulty related to the displacement of the sheets without fracturing of the lithological units resulting in a complete collapse during deformation, and he pointed out that johnstrup (1882) had suggested that the deformed layers could have been ground-frozen. jessen’s (1931) more subjective arguments against slater’s work concern the fact that slater (1927) did not refer to jessen’s (1918) substantial work on vendsyssel and in particular his published cross-section of lønstrup klint. jessen pointed out that major anticlines in slater’s crosssection between mårup kirke and rubjerg knude fyr do not exist, and that slater’s (1927) misinterpretation must be ascribed to his superficial investigations which did not allow him to check the way-up relationship of each limb in the fold structure (jessen 1931). in his work on the glaciotectonic deformation of palaeogene diatomites with ash layers in the limfjorden region, gry (1940) compared these with the deformation at lønstrup klint and supported the proglacial deformation concept of gripp (1929) and jessen (1931). furthermore, gry proposed a gravity-spreading model for the deformation and attempted a very early balanced cross-section in the consideration of restoration of the dislocated thrust sheets (fig. 3). however, gry (1940) proposed a cylindrical model for the thrust surfaces, and in his ‘back-stripping’ cross-section the floes were displaced along circular fault lines. thus, in his dynamic consideration the floes were assigned a standing position with their frontal parts ‘up in the air’ (fig. 3), and he consequently concluded that more than 80% of the upper sand-series at lønstrup had been eroded away by the advancing ice. in contrast to this point of view, pedersen (1987) suggested that a large proportion of the upper sandseries was deposited syntectonically; this removed the requirement that a large part of the floes or thrust sheets had been eroded away. pedersen (1987) interpreted the glaciotectonic thrust-fault complex as an example of gravity-spreading deformation, viewed in the light of the gravity-spreading experimental model presented by bucher (1956) and with reference to comparable gravity-spreading deformation in soft sedimentary rocks exemplified by the mudlumps in the mississippi delta (morgan et al. 1968). furthermore, the mudlumps or mud diapirs in the lønstrup klint imbricate fan were described, and interpreted as an integral part of a conceptual dynamic model for thrust-fault related mud diapirism and syntectonic sedimentation (fig. 4). sadolin et al. (1997) elaborated on the model of syntectonic sedimentation in the lønstrup klint section. based on detailed sedimentological studies, they pointed out the importance of the unconformity that separates the lower muddy units (their unit a), from diluvial sand yoldia clay fig. 3. a model for structural balancing of the dislocated floes in the lønstrup klint section suggested by gry (1941). in his model, the displacement surfaces were regarded as cylindrical sections and due to the suggested amount of displacement about 80% of the dislocated floes was subsequently eroded away. 16 the upper sandy units (their units b–d). the lower unit a was interpreted to have been deposited in a lake isolated from the former marine kattegat–skagerrak basin by either a damming of the advancing ice, in accordance with ideas also presented by jessen (1918, 1931), or simply by isolation of the lake basin due to lowering of sea level in the late pleistocene (sadolin et al. 1997). the unconformity was interpreted to reflect a major drainage event of the lake basin before a shallow lacustrine basin was established, characterised by incursions of glaciofluvial deposition (units b–d of sadolin et al.1997). during the deposition of units c and d, glaciotectonic thrusting commenced contemporaneously with the rise of mud diapirs and the formation of normal faults due to mass adjustments in the mobilised mud in the subsurface (sadolin et al. 1997; fig. 5). the conceptual model presented here aims at an interpretation based on the concepts of thin-skinned thrust-fault tectonics. although the scale is an order of magnitude smaller than in typical orogenic belts, it has not been found appropriate to introduce special terminology for the deformation structures in the rubjerg knude glaciotectonic complex. the concept of thrust-fault deformation and related structures is summarised in the following chapter. thin-skinned thrust faulting: the concept it is difficult to judge exactly when the concept of thin-skinned thrust faulting nucleated, as it represents a gradual evolution of ideas over the last 25 years or more. however, boyer & elliot (1982) appear to have been the first to give a conceptual introduction to the basic principle of thin-skinned thrust faulting. suppe (1983, 1985) improved the concept by defining and describing the geometry and kinematics of fault-bend folding. jamison (1987) and schirmer (1988) contributed with further improvements of geometric analysis of fold development in overthrust terranes and thrust-fault hanging-wall successions. mcclay (1992) presented a glossary of thrust tectonic terms, and erickson & jamison (1995) demonstrated viscous-plastic finite-element models of fault-bend folds. in 1997, an entire volume of the journal of structural geology was devoted to thrust-fault tectonics. among the papers that particularly inspired and supported this study of glaciotectonic thrust faulting were those of contreras & sutter (1997), medwedeff & suppe (1997) and mitra & sussman (1997). thrust-fault modelling to better understand the range of possible configurations of different structural frameworks of thrust-fault complexes, a series of computer models were tested with the aid of the program autofault, a ‘balanced cross section program’ within the autocad system frame (ozkaya 1994). four of these test models are demonstrated here to illustrate the thin-skinned thrustfault concept (figs 6–9). the basic function of the model is to define and construct a layer package onto which a thrust fault is added and given a certain displacement. the program then calculates the configuration of the thrust sheet fig. 4. a four-stage model for the development of mud diapirs related to thrust faulting in lønstrup klint suggested by pedersen (1987). note that in the model the thrust zone of the hanging-wall ramp constitutes mobilised mud and that syntectonic deposits accumulate ‘piggyback’ between the thrust sheets. 17 100 m 100 m 100 m 100 m sandrende thrust fault c c c b b b b a a a a a b c a mud diapir x1 a x2 a c c mobilized mud y1 y2 extensional faults mobilized mud x1 x2 y1 y2 x1 x2 y1 y2 position of thrust-fault fractures unconformityerosionally removed 2 1 d d m. s. l. older yoldia clay 4 3 fig. 5. the structural and depositional development of the sandrende section suggested by sadolin et al. (1997). the model summarises four stages of development initiating with the formation of the regional erosional unconformity (1). unit b was deposited in topographic lows above the unconformity, and thrust faulting initiated contemporaneously with the deposition of unit c (x 1 –y 1 and x 2 –y 2 denote same reference points separated by the thrusts, where x = footwall syncline and y = hanging-wall anticline) (2). propagation along the thrust faults continued and unit c was deposited during increasing tilting of the thrust sheet. normal-fault fractures formed in connection with the incipient diapirism (3). the sandrende diapir rose during deposition of unit d and normal faulting propagated. in the proximal part of the thrust sheet, a network of conjugate extensional faults developed and interference between a new-formed satellite thrust and the normal faults af fected the complex. the tip of the thrust sheet was bent due to drag along the side of the rising diapir (4). star symbol provides a reference point through the development stages. 18 step 1 step 2 step 3 step 4 step 5 step 6 hanging-wall block ramp footwall block 50 m displacement 100 m displacement 150 m displacement 200 m displacement 300 m displacement 400 m displacement lower flat axial surface upper flat = top surface lower ramp hinge upper ramp hinge hanging-wall anticline hinterland-dipping limb foreland-dipping limb hanging-wall flat hanging-wall flat hanging-wall ramp upper footwall flat footwall ra mp décollement or lower footwall flat 0 100 200 300 400 m n s fig. 6. test model 1 of thrust-fault deformation constructed with the computer program autofault (ozkaya 1994). the model demonstrates the development of simple ramp propagation given increasing displacements. in the first four steps, the displacement is sequentially increased by 50 m, whereas a displacement of 100 m is added to steps 5 and 6. note that a ‘typical upright anticline’ develops when the displacement is about twice the thickness of the layer package displaced. moreover, the model illustrates the terminology applied in the text and defined in appendix 1. 19 for the specific model constructed. thus the program gives the ‘differential’ calculation model to an induced ‘integration’ solution configuration. further thrust faults can be added, and be given new displacements, such that rather complex models can be constructed. however, a few limitations of the program hamper realistic comparisons with nature. thus the program cannot handle inclinations exceeding 60°. in general this is not a problem as ramp angles typically range between 10° and 35° and for rock mechanical reasons never exceed 45° (ozkaya 1994). however, the problem of steep inclinations becomes important in complexes including superimposed deformation. a second limitation is that testing with superimposed displacements requires a construction with an upper flat located within the model. this results in an unrealistically high number of shallow upper flats in the models, as illustrated below in test model 4 (see fig. 9). thirdly, the program cannot accommodate cross-cutting thrust-fault relationships, which limits the spacing and dip of ramps. nevertheless, the test models give a good introduction to the thrust-fault concept, and demonstration of models with basic layer package dimensions approaching the scale of thrust sheets involved in the rubjerg knude glaciotectonic complex can be achieved. a glossary of the thrust-fault terms used here is given in appendix 1; note that only contractional thrust-fault structures are considered. test model 1 the first autofault model displays a simple thrust fault with one ramp connecting a lower and an upper flat (fig. 6). the development of thrust-fault structures, in particular the fault-bend folding of the hangingwall anticline, is demonstrated in six steps with increasing displacement. the ramp angle is 25°, and the layer package constitutes a lower unit 25 m thick where the lower flat (or the décollement zone) is located. above this, one 25 m and two 20 m thick layers have been constructed, with a 30 m thick uppermost layer (fig. 6). the model approaches the assumptions of parallel behaviour with preservation of layer thickness, no net distortion where layers are horizontal, and conservation of bed length (suppe 1983). step 1 shows the gentle hanging-wall anticlinal folding after 50 m displacement. note the flat-topped nature of the hanging-wall anticline, which makes it almost insignificant. the backlimb of the anticline dips toward the left, parallel to the ramp, and the axial surfaces defined by the bend above the lower ramp hinge and the bend of the hanging-wall anticline define two kink bands dipping steeply to the right. by comparing steps 1 and 2 it can be seen that the spacing between the kink bands increases with increasing displacement. step 2 gives the configuration after 100 m displacement. here the forelimb dipping towards the foreland to the right starts to be a significant part of the structure. note the increase in spacing between the kink bands in the backlimb structure. the kink bands define minor zones of weakness, which could develop into small reverse faults as in the thrust model demonstrated by wiltschko (1979). these are referred to as back thrusts. step 3 shows the structural development after 150 m displacement. note that the flat-topped hangingwall anticline now has a more angular upright form, where the kink bands fanning up from the positions near the upper ramp hinge approach each other. however, in the model the anticline maintains its flat-topped structure and retains two axial surfaces (kink bands). step 4 demonstrates the formation of the upright, angular hanging-wall anticline, where the amount of displacement is close to the length of the thrust-fault ramp. due to the geometric adjustments the hangingwall ramp is shorter than the footwall ramp. the displacement is 200 m corresponding to about two times the thickness of the thrust sheet. step 5 shows the structural development after 300 m displacement. the hanging-wall anticline becomes even more flat-topped and the space between its axial surface kink bands increases. note that the forelanddipping forelimb is linked to the hanging-wall ramp displaced along the footwall flat, and the hinterlanddipping backlimb corresponds to the hanging-wall flat bent up along the footwall ramp. step 6, with a displacement of 400 m demonstrates that the main structural configuration is maintained, except for the increase in spacing between the backlimb and the forelimb. test model 2 the second autofault model demonstrates the propagation along a thrust fault dif ferentiated into a décollement zone, a lower ramp, an intermediate flat, an upper ramp and an upper flat bringing the thrust fault up to the top surface (fig. 7). the model is con20 structed with two lower units, 40 m in thickness; the décollement zone is located in the second layer. the lower layers mimic the lower clay units of the lønstrup klint stratigraphy, and two c. 25 m thick layers overlie them. the top layer is 50 m thick, but while not comparable to any part of the stratigraphy in the lønstrup klint section, its construction yields a better demonstration of the development envisaged. the lower ramp is given a dip of 25° and the upper ramp a dip of only 15° to reflect the principle of increasing angle of fracturing with increasing depth (hobbs et al. 1976; pedersen 1996). the distance between lower and upper ramps along the intermediate flat is c. 250 m, and three steps are presented in fig. 7. step 1 is given 50 m displacement and two hanging-wall anticlines immediately appear. the steep ramp clearly initiates the formation of an upright anticline with steeply dipping limbs. between the two hanging-wall anticlines, an intervening syncline forms above the intermediate flat. the involute surface of the syncline provides the location for a broad, shallow basin. step 2 shows the structural development after 100 m displacement. this demonstrates clearly that the intervening syncline is an obvious site for a piggyback basin to develop. note that the steeply dipping forelimb of the hanging-wall anticline above the lower ramp would be the obvious site for erosion and the source of material feeding into the piggyback basin. step 3 demonstrates that with a displacement of 200 m, the piggyback basin becomes narrow and is elevated to a higher position as a consequence of the displacement up along the upper ramp; it is eventually lifted out of the position for being a centre of deposition. with increasing displacement, the frontal part of the thrust sheet develops into a wedge-shape structure. hanging-wall block lower hanging-wall ramp lower ramp upper ramp upper flat upper hanging-wall ramp upper hanging-wall ramp footwall block 50 m displacement 100 m displacement 200 m displacement piggyback basin piggyback basin intermediate flat lower flat fault-bend folding step 1 step 2 step 3 0 100 200 m n s fig. 7. test model 2 of thrust-fault deformation constructed with the computer program autofault. the model demonstrates the development of thrust-fault propagation along a lower and an upper ramp and the connecting flats. note in this model the formation of two anticlines divided by a syncline, the depression of which is the obvious location of a piggyback basin. 21 test model 3 the third autofault model aims at constructing an imbricate complex by branching faults fanning up from the same décollement level (fig. 8). the model is constructed with a lower 20 m thick unit in the top of which the décollement zone is located. above the décollement zone, three units with a combined thickness of 50 m form the lower part of the thrust sheets, and the succession is capped by an upper 20 m thick unit. in three sequential steps, the principle of piggyback thrusting is demonstrated (fig. 8). step 1 shows 100 m displacement along a deeprooted ramp dipping 30°. note the normal architecture of the hanging-wall anticline results from the ramping (compare with fig. 6, step 3). step 2 demonstrates the re-orientation of the piggyback thrust sheet by the introduction of 100 m displacement along a 18° dipping ramp in front of and below the first thrust fault. note that the accumulated displacement of the first thrust sheet amounts to c. 200 m. step 3 shows an additional 100 m displacement along a low-angle 12° dipping ramp. although the model demonstrates the main architecture of the imbricate fan illustrated by pedersen (1987), it is a fairly simple model which may have only little relevance to natural conditions. test model 4 the final autofault model demonstrates the more likely formation of a steeply dipping imbricate fan or duplex (fig. 9). the model is given the same stratigraphic units as in test model 3 (fig. 8). a longer décollement zone is located in the middle of the lowermost unit, in addition to an intermediate flat in the third layer, while the upper flats are located within the uppermost unit. the initial steps in the construc1 1 2 3 1 2 100 m displacement 100 + 100 m displacement 300 m accumulated displacement simple ramp piggyback thrust sheet branching thrust fault branching thrust-fault imbricate fan step 1 step 2 step 3 0 100 200 m n s fig. 8. test model 3 of thrust-fault deformation constructed with the computer program autofault. the model demonstrates the formation of an imbricate fan by successive thrust-fault splays branching up from the main décollement zone. the encircled numbers refer to the sequential phase of thrust imbrication. the model is probably not comparable to structures formed in nature, but can be regarded as an introduction to test model 4 (fig. 9). 22 tion of this model are similar to the examples demonstrated above, and hence only the final two steps are illustrated (fig. 9). however, these give a convincing illustration of the increase of dips in an imbricate thrustfault complex. step 1 illustrates the final structural architecture after 140 m displacement of thrust sheet 1 along the décollement zone, the lower ramp, the intermediate flat, an upper ramp and onto the upper flat (dips of ramps c. 25°). thrust sheets 2–5 were formed by branching ramps (dip of ramps c. 15°) with a displacement of c. 80 m added to each thrust fault. finally, the leading thrust sheet (6) is displaced 90 m along the lower décollement zone and a deep-rooted 30° dipping ramp. note that the branching ramp imbricates are carried piggyback on thrust sheet 6. furthermore, it should be noted that a long trailing segment of thrust sheet 6 occurs between the décollement zone and the intermediate flat. if this trailing segment becomes chopped up into duplexes between the two deep-rooted ramps, it will affect the overlying imbricates by vertical elevation and the formation of antiformal stacks. step 2 illustrates the over-steepening of the imbricates stacked onto the backlimb of the hanging-wall anticline of thrust sheet 6 arising from the addition of 100 m displacement to step 1 along the leading thrust rooting down to the lower décollement zone. test models: concluding remarks a set of principles may be derived from the test models. 1. the level of elevation of the reference surface is directly related to the number and sizes of ramps the thrust sheet has passed. a ramp rooting down to a deep flat level corresponds to a high elevation of the topmost reference surface. in contrast, if a top reference surface is positioned at the same level as in the foreland, the thrusting corresponds to a translation along a flat. 2. the steeper the ramp, the earlier its time of formation. gently dipping ramps are initiated at a late stage of deformation in areas proximal to the foreland. 3. the thickness of a piggyback basin reflects its duration as depocentre. thus a small thickness of piggyback basin fill indicates an early trapping of the basin by overthrusting of a hanging-wall block. 4. a thick succession in the piggyback basin reflects a long period of translation of the thrust sheet along a long flat. step 1 step 2 fig. 9. test model 4 of thrustfault deformation constructed with the computer program autofault. the model demonstrates an imbricate fan (see fig. 8) subjected to faultbend folding during piggyback translation of an underlying hanging-wall flat propagation along a footwall ramp. the footwall ramp propagation will consequently result in increasing dips of the thrust sheets in the imbricate fan. encircled numbers indicate successive thrust sheets. concept of balanced cross-section the principle of the balanced cross-section in structural analysis of thrust-fault systems was elegantly outlined by dahlström (1969) and further improved by suppe (1985). the application of balanced crosssections in glaciotectonics has been demonstrated by croot (1987), klint & pedersen (1995) and pedersen (1996). in the construction of the balanced section, two different functions are applied: (1) the line balance, and (2) the volume balance, which in a 2-d crosssection corresponds to area balance. the first function concerns the length of displacement, whereas the second function concerns the preservation of volume in the deformed cross-section compared with the restored undeformed cross-section (for demonstration see plate 2). the basic method of balancing a crosssection (dahlström 1969) is restoration by defining a pinpoint to be fixed to the foreland and then restoring the thrust sheets back to their initial pre-deformational position. thus one begins at the foreland and then by line balancing the thrust sheets are pulled back sequentially to their position prior to displacement. this requires a measure of displacement, which is the essential, but often difficult figure to achieve without some range of uncertainty. details concerning the construction of the balanced cross-section of the rubjerg knude glaciotectonic complex (plate 2) are given below. 23 location the rubjerg knude cross-section is 6124 m long and extends from the coastal cliff immediately south of lønstrup, ribjerg, to about 300 m north of the ramp leading down to the beach at nørre lyngby (fig. 2, plate 1). the strike of the section is 17°, which is nearly parallel to the direction of the coastline. this is also approximately perpendicular to the main concentration of structural strikes (bedding, thrust faults and fold axes; fig. 10). the cross-section was consequently constructed to fit a general plane of orthographic projection with a projection axis striking 107°. the rubjerg knude cross-section covers only the rubjerg knude glaciotectonic complex. thus it is not as extensive as the cross-section of lønstrup klint constructed by jessen (1918, 1931), which extends from the cliff at the northern fringe of lønstrup to the northern part of the beach at løkken (see fig. 12). the utm co-ordinates (zone 32, ed50) of the end points of the rubjerg knude cross-section are 547512, 6370243 (n-end point) and 545251, 6364783 (s-end point). photogrammetric work the cross-section of rubjerg knude glaciotectonic complex (plate 1) is based on a multi-model photogrammetric investigation of the cliff section using the method described by dueholm (1992). oblique photographs were taken from a cessna fixed-wing aircraft with a minolta xg2 camera with known optical specifications, calibrated at the laboratory of photogrammetry at the technical university of denmark. standard 24 × 36 mm diapositive colour film was used, and the photographs were taken with 66% overlap from a distance of 200–300 m with an inclination angle of c. 35°, which provided the basis for setting up 67 stereoscopic models. in the laboratory, the orientation of the stereo-models was carried out based on ground control points adapted from two sets of vertical aerial photographs at a scale of 1:25 000, namely d9202 g 1365–66 and kms 9203 a509–10 taken in may 1992. the stereoscopic instrument used was a kern dsr 15 analytic plotter with a dec vms operating system and the special attached geoprogram developed n s 0 100 200 m 100 m displacement on youngest thrust fault 200 m displacement on youngest thrust fault 1 1 2 3 5 2 3 4 5 4 6 6 location and construction of the rubjerg knude cross-section 24 1 2 3 n n = 52 1 2 3 n = 60 n 1 2 3 n = 83 n 1 2 3 n = 83 n a b c d fig. 10. stereographic projection diagrams of the orientation of structural elements in the rubjerg knude cross-section. the stereograms, lower hemisphere, equal area (schmidt) net, display the concentration of the poles to bedding planes (black dots) or thrust planes (black triangles). a and b are measurements taken from jessen (1931), and c and d are data produced in this study. contour intervals are 1, 2.5, 5, 7.5, 10, 12.5, and 15%. the density point in all four diagrams is close to 197°/35°. comparing the two sets of diagrams demonstrates that the structural orientation has been maintained despite c. 100 years erosion corresponding to c. 125 m retreat of the coastal cliff section. black squares (d) indicate normal fault planes. blue lines/numbers indicate principal compression axes. 25 by dueholm (1992). in the stereoscopic models, the geological features were outlined by the floating mark and digitised by the attached computer. the digitised data were stored for the later construction of the crosssection and the transformation for other programs applied for the management of the cross-section display. the scale of the rubjerg knude cross-section in the analytic plotter version is 1:500, and the accuracy of the plotted data is about 25 cm (for further details, see appendix 2). digital editing in order to represent the cross-section in a publishable display, the digitised data were transferred to arcinfo at the gis-laboratory at the geological survey. here it was transformed into an arc-view project, which served as the computer tool for editing the crosssection. thus all areas were converted to closed polygons, which were annotated to fit the legend of lithologies. during this editing, interpretations were made to finish the display of the cross-section, in particular interpretations of the scree-covered parts of the section. this was carried out contemporaneously with the construction of the balanced cross-section (see below), and a few additional corrections were added to the rubjerg knude cross-section. some new exposures along the cliff section appeared in 1997–1999, which added to a better understanding of the structures in the transition from the frontal part of the glaciotectonic complex to its foreland. these have been incorporated into the arc-view project. the final editing of the cross-section concerned the balanced cross-section. the construction of the balanced section was digitised and transformed into an arc-view project, and the subsequent interpretation of the extension of the thrust-fault ramps below sea level was added. thus the rubjerg knude cross-section comprises a display of the exposed part of the cliff section with lithological and structural identity added as themes. furthermore, the cross-section includes an interpretation of the thrust-fault structures in the subsurface. finally, a balanced construction was added l l 2900 3000 m dc dm ds α α l/r-u l/r-u t t t t l/r-u l/r-u fig. 11. illustration of the method used for estimation of the displacement for the balanced cross-section. above the main erosional unconformity at the top of the cliff, the extension of the thrust sheet tip is constructed by the intersection between the thrust fault (t) and the l/r-unconformity (l/r-u) based on the angle (±) between the bedding of the thrust sheet and the hanging-wall ramp. dm, displacement measured; dc, displacement constructed from tip-extension; ds , displacement estimated from the interpretation of thrust-fault trace under the scree cover. the section illustrated is part of the rubjerg knude fyr section (plate 1). 26 table 1. the distribution of areas in the balanced cross-section (plate 2) balance (plate 2a) ramps (plate 2b) section* number of areas area (m2) section* number of areas area (m2) 01ul 5 23 048 01ul 13 24 302 02sn 13 8965 02sn 18 8536 03mb 15 28 443 03mb 21 30 944 04kr 10 24 158 04kr 22 18 390 05br 28 34 143 05br 40 33 548 06sr 28 33 218 06sr 49 31 588 07ss 32 26 421 07ss 31 23 973 08gr 55 49 842 08gr 47 45 118 09rf 30 22 827 09rf 26 21 458 10st 54 43 674 10st 41 36 656 11mr 69 51 902 11mr 55 45 342 12mk 95 82 226 12mk 87 62 763 13bl 8 17 922 13bl 2 14 313 nrly 2 5437 13ri 1 4405 ptr 3 2538 md 1 472 ve 4 9818 * the annotated numbers of sections (05br) correspond to the sequential location of each section in a distal–proximal order, and the capitalised letters refer to the general abbreviation of the section names (see plate 2). to the cross-section project, such that each thrust sheet is annotated in a balanced restored cross-section as well as in the structural cross-section displaying the geometry of the ramps and flats (plate 2). construction of the balanced cross-section the construction of the balanced cross-section for the rubjerg knude glaciotectonic complex was based on the geological cross-section, which displays the geometry of the thrust sheets in sufficient detail to allow calculations of their displacements and cross-sectional areas (plates 1, 2). the method of balancing necessitates that the thrust sheet closest to the foreland is the first to be restored to its pre-deformational position. therefore, the balancing works backwards from the distal to the proximal deformation area, and consequently the annotation of the thrust sheets begins with the first thrust sheet restored. in the balanced cross-section of the rubjerg knude glaciotectonic complex, the thrust sheets are additionally annotated according to that part of the cliff in which they occur: two capital letters refer to the name of the section and a number refers to its position from leading edge to trailing end of the section. thus, kr01 is the thrust sheet nearest to the foreland in the kramrende section. a thrust fault is referred to according to the thrust sheet it displaces. however, the trailing footwall ramp is referred to the annotation of the footwall block, which underlies the hanging-wall ramp/flat of the thrust sheet displaced over it. thus the kr02 hanging-wall ramp is displaced up along the kr01 footwall ramp. although one of the basic conditions in constructing balanced sections is the preservation of volumes, which in the areas strongly affected by mud remobili27 sation and diapirism is difficult to maintain, the exercise has been carried out to match a balanced section to the mapped and interpreted thrust-fault framework. so despite the uncertainties and the demand for interpretation of the geometry and magnitude of eroded thrust sheet tapers, the construction of the balanced section added significantly to the understanding of the duplex framework (plate 2b). in the rubjerg knude cross-section (plate 1), the displacement is measured and estimated mainly from the distance between the intersection of the l/r-unconformity (the unconformity between the lønstrup klint and rubjerg knude formations, see below) and the footwall ramp, and the intersection of the l/runconformity and the hanging-wall ramp (fig. 11). however, the tips of the thrust sheets are generally eroded away, so the first approximation is from the l/r-unconformity footwall point to the point where the hanging-wall ramp is truncated by the glaciotectonic unconformity at the top of the cliff. the second approximation is the addition of the distance estimated from the size of the tip eroded away. this estimate is based on a simple geometric construction of the tip-triangle from the dips of the hanging-wall ramp and the l/runconformity, respectively (fig. 11). this line balance is subsequently controlled by the width of the piggyback basin more or less corresponding to the upper footwall flat. all the measured displacements are strictly restricted to the minimum distance to avoid unrealistic exaggerations. therefore the actual displacements might be slightly greater. the area balance is based on a calculation of all the areas annotated in plate 2. the computer-supported calculation was carried out with the arc-info program, and the calculations of the areas in the balanced cross-section and the ramp cross-section deviate by less than 10% (plate 2a, b). this is regarded as a reasonable correspondence considering the various sources of error (table 1). in general, the sections have a smaller area in the ramp cross-section (plate 2b) due to the erosion of areas above the main headof-cliff unconformity, and in most sections the number of areas is higher due to the increased complexity of the geometry in the reconstructed structural cross-section (plate 2b). geological setting the rubjerg knude glaciotectonic complex incorporates deformed sedimentary deposits that belong to the upper part of the mainly marine succession known previously as the skærumhede series (jessen et al. 1910). this succession was deposited in the northern part of the danish basin in the late pleistocene, after the late saalian terrestrial glaciation retreated from denmark (houmark-nielsen 1987, 1999; knudsen 1994). the major source area for deposits in this part of the danish basin is the scandinavian basement in southern norway and central sweden, that comprises precambrian fennoscandian granites and gneisses overlain by palaeozoic metasediments, including permian volcanics and their related intrusive magmatic rocks of the oslo province (oftedahl 1981). the extrabasinal indicator boulders reflect these source areas, which were situated between the centres of ice-cap nucleation and the depositional basin (milthers 1909; smed 1995). the boundary between the northern part of the danish basin and the south-western part of the elevated scandinavian basement is covered by the skagerrak, the sea covering a deep depression (about 500 m deep) known as the norwegian channel (sejrup et al. 1987, 1994, 1998). one of the important discussions concerning the glaciation of denmark during the last stadial focuses on how the ice from norway advanced across the skagerrak about 30 000 years ago. the problem involves the dynamics of the ice stream along the southern coast of norway, the so-called norwegian channel ice stream, and the interaction between the marine and terrestrial parts of the ice cap in southwest norway (larsen et al. 2000). associated problems include the filling of the deep trench in skagerrak, and the termination of marine conditions in skagerrak, vendsyssel, and the northern north sea as well as the kattegat (for locations, see fig. 12). the marine environment referred to as the older yoldia sea, which extended into the vendsyssel region, formed in the late saalian, and the climatic change from a mild climate in the eemian to a glacial 28 ab 58˚n 58˚n 56˚n 56˚n skagen hirtshals frederikshavn mols djursland samsø hven hundested glumsløv ristinge klint bovbjerg jylland jylland læsø anholt langeland 0 100 km50 norway sweden limfjorden skagerrak north sea skagerrak kattegat kattegat vendsyssel mors germanyfig. 13 lodbjerg vendsyssel sd kt msl m sl lim fj or d en i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i denmark 12˚e 12˚e lønstrup lim fjorden jæren n orwegia n tren ch fyn sjælland skåne lønstrup klint 29 climate in the weichselian is recorded in a series of wells drilled in north jylland and the kattegat region (knudsen & lykke-andersen 1982; lykke-andersen 1987; lykke-andersen & knudsen 1991; knudsen 1994). towards the end of the middle weichselian the scandinavian ice sheet over southern norway built up. the ice streams were drained from a main spillway in oslo fjord moving out through the norwegian channel along the coastline of southern norway (larsen et al. 2000). a change in the dynamics of the scandinavian ice sheet over southern norway forced the glaciers to progress south-westward across the norwegian channel. the ice advanced into the northern north sea, where a glacial cover was established about 29 000 years b.p. and lasted until 22 000 years b.p., when the first recurrence of marine conditions (the ‘young yoldia sea’) was recorded (sejrup et al. 1994, 2000). this glacial coverage was probably closely connected with the fall in sea level, amounting to 120 m below present sea level (fairbanks 1989; bard et al . 1993), which could have hampered the active drainage of the norwegian channel ice stream. the ice spread southward over the skagerrak causing the kattegat basin to be dammed by the ice margin and terrestrial areas to be established in the central part of the north sea (sadolin et al. 1997; houmark-nielsen 1999). as a consequence, the kattegat–skagerrak region began to dry up due to the general sea-level fall; this is reflected in the progression from arctic marine conditions in the skærumhede series to brackish and glaciolacustrine environments. this change took place at about 32 000 years b.p. (table 2), and may have been accentuated by the addition of meltwater from the advancing norwegian ice (jessen 1918; sadolin et al. 1997). the dramatic drainage of the lake basin in the kattegat towards the north sea is recorded by a significant erosional unconformity in the sedimentary succession at lønstrup klint (the l/r-unconformity), dated as close to 29 000 years b.p. (sadolin et al. 1997). shortly afterwards, the basin was once again dammed and shallow lacustrine and fluvial environments were established while proglacial thrust faulting was initiated reflecting the relatively fast advance of the ice margin (sadolin et al. 1997). the thin-skinned thrust faulting in the rubjerg knude glaciotectonic complex involved an accretionary wedge extending more than 12 km to the south in front of the advancing ice margin. the lowermost décollement level was situated in the marine clays of the older yoldia sea. after a compression of about 50%, the glaciotectonic complex was formed (pedersen 1987) leaving a large part of the area between lønstrup and hirtshals as a depression corresponding to the ‘hole’ and the rubjerg knude glaciotectonic complex to the ‘hill’, in a ‘hill-andhole’ pair in the sense of aber et al. (1989). subsequently the norwegian ice truncated the glaciotectonic complex and the deposition of the kattegat till formation concealed its structures. the norwegian ice advanced down to a stationary line (figs 1, 12) crossing central denmark from west to east, whose position is inferred from the distribution of the kattegat till formation (houmark-nielsen 1987, 1999, 2003; pedersen & petersen 1997). after its termination at the stationary line (figs 1, 12), the norwegian ice melted back. it was succeeded by the main south-west ice advance of the scandinavian ice sheet, which extended out to the main stationary line (ussing 1903; houmark-nielsen 1987, 2003; pedersen et al. 1988). in northern jylland, the isostatic depression due to the loading of the ice sheet was substantial. the termination of the glaciation in denmark thus resulted in interference between eustatic sea-level rise and isostatic rebound with a complex depositional development during the re-establishment of the younger yoldia sea in the skagerrak– vendsyssel–kattegat region about 17 000 years ago. this may be summarised as a forced regression under progressively falling sea level due to the isostatic rise of the vendsyssel region (richard 1996). the vennebjerg and rubjerg knude hilly islands probably formed part of a larger island archipelago extending out into the north sea. terrestrial conditions were established at the end of the weichselian. at nørre lyngby (fig. 13), a depression was formed above a neotectonic fault zone that predated older dryas time (lykke-andersen 1992). in this depression, lacustrine gyttja and fluvial sand of older dryas and allerød age were deposited; a large number of mammalian remains have been found in these deposits indicating an arctic to sub-arctic reinfacing page: fig. 12. location map. map (a) shows the main part of the danish basin with the surrounding land areas. sdkt is the position of the stationary line for the norwegian ice advance (sdkt is an abbreviation of southern distribution of kattegat till fm). msl is the main stationary line for the scandinavian ice sheet at the glacial maximum in the late weichselian. map (b) gives the position of relevant geographical localities in denmark as well as the location of fig. 13, the geological map of vendsyssel. 30 deer steppe also populated by hunters (jessen & nordmann 1915; aaris-sørensen 1995). during holocene time, the vendsyssel region was affected by isostatic rebound (mertz 1924). at lønstrup klint, this resulted in a 25 m elevation of the heterolithic sediments of the younger yoldia sea. bogs developed in the depressions on the glacial peneplain at the end of the stone age and the beginning of the bronze age (jessen 1918). up to 1.5 m of peat accumulated; when this is exposed in the clif f surface and blocks of peat fall down onto the beach, the peat is locally called martørv (sea-peat). the locality names martørv bakker (sea-peat hill) and moserende (boggully) refer to these deposits. the geomorphology of the cliff is strongly influenced by the thrust-fault structures. the clayey parts of the thrust sheets form ridges that form projections along the coast between gullies that are eroded out in the sandy parts (schou 1949). springs typically well out at the surface between the clayey and sandy units and more incised gullies (render in danish) are formed where the drainage is concentrated. although the location of gullies and the clif f line have retreated about 100 m since a. jessen constructed the first cross-section of lønstrup klint, it has been possible to retain his names in the present cross-section (plate 1). the general erosion rate of the cliff is about 1.3–1.5 m per year (jessen 1918; pedersen 1986b). landslides occur very frequently, particularly at sites where mud diapirs are located in the cliff section. where glaciofluvial deposits dominate the cliff section, there is a marked tendency for aeolian dunes to accumulate on top of the cliff (pedersen 1986b). wind action on the moraine plateau on top of the cliff has eroded the finegrained material away from the till deposits, leaving a stone pavement as the residual trace of the glacially truncated surface. aeolian sand migration intensified about 300–400 years ago (jessen 1918), one of the consequences being the burial and abandonment of the old rubjerg church. the high aeolian dunes on top of rubjerg knude have accumulated during the last 100 years. the rubjerg knude lighthouse was built in 1900 (bendsen 1981) when dunes were less than 10 m high. today the tops of the dunes are close to 100 m above sea level corresponding to a vertical dune accumulation of nearly 50 m. the present-day steep nature of the dunes was probably stimulated by the artificial dune protection fences. however, the steady erosion of the cliff indicates that the lighthouse will fall into the sea about ten years from now. table 2. radiocarbon dates, rubjerg knude and lønstrup klint, vendsyssel, northern denmark stratigraphic unit vendsyssel fm vendsyssel fm vendsyssel fm rubjerg knude fm rubjerg knude fm lønstrup klint fm stortorn fm stortorn fm stortorn fm locality lønstrup klint lønstrup klint lønstrup klint sandrende lønstrup klint sandrende ribjerg mårup kirke stortorn lab. id no. k-858 k-2670 aar-2134 aar-2265 aar-4066 ua-4454 aar-4067 aar-4068 aar-4069 material mollusc mollusc mollusc plant mollusc moss mollusc mollusc mollusc 14c age ka b.p. 13.9 ± 0.2 14.7 ± 0.2 14.5 ± 0.2 30.9 ± 0.5 43.0 ± 1.3 29.2 ± 1.4 29.6 ± 0.4 30.9 ± 0.4 31.3 ± 0.4 calib. age ka b.p.* 16 ± 1 17 ± 1 17 ± 1 33 ± 1 46 ± 3 32 ± 1 33 ± 1 34 ± 1 34 ± 1 13c‰ pdb+ 0.6 –27.3 3.3 –29.1 1.5 1.7 1.3 ref.‡ (1) (2) (3) (4) (5) (4) (5) (5) (5) * calibrated ages are calculated according to bard et al. 1993 and kitagawa & van der plicht (1998). + relative to pdb standard.‡ references: 1: krog & tauber (1974); 2: knudsen (1978); 3: richardt (1996); 4: houmark-nielsen et al. (1996); 5: this study. 31 holocene vendsyssel formation glacigene formations upper cretaceous chalk 20 km5 10 150 10˚30’10˚00’ nørre lyngby mårup løkken rubjerg knude skærumhede skagen frederikshavn hirtshals kattegat skagerrak location of well lø ns tr up k lin t lønstrup sandrende vennebjerg dgu no. 1.287 dgu no. 10.4 dgu no. 8.137 stortorn 57˚30’ dgu no. 8.137 fig. 13. geological map of vendsyssel showing the location of three wells referred to in the text. 32 h ol oc en e m id dl e w ei ch se lia n la te w ei ch se lia n sk æ ru m he de g ro up 11.5 19–15 27–23 30–28 32–30 35–32 fo rm at io n u nc on fo rm ity li th ol og y c la y si lt pe bb le sa nd c ob bl e fo ss ils st ru ct ur es an d gr ai n si ze s d ir ec tio na l el em en ts t hi ck ne ss a ge k a b. p. (c al en da r ye ar s) 30 20 10 0 m recent aeolian dunes main cliff-section unconformity nr. lyngby beds vendsyssel formation mid danish till formation ribjerg formation blå-unconformityeu kattegat till formation rubjerg knude formation eu l/r-unconformity lønstrup klint formation stortorn formation gu gu eu eu clay silty mud dropstones in clay and silty mud sandy mud sand till gravel plant fossil dated shell dated palaeocurrent direction of ice movement erosional unconformity glaciotectonic unconformity f. m. c. eu gu peat and gyttja 33 lithostratigraphy the upper pleistocene in the vendsyssel region comprises three major stratigraphic units: (1) a unit of marine sediments deposited on the erosional surface of the saalian till, overlain by (2) a glacioterrestrial succession that in turn is succeeded by (3) a second marine deposit. the first marine unit was laid down after the retreat of the ice from the region at the end of saalian time. during eemian and early to middle weichselian time, the older yoldia sea prevailed (jessen et al. 1910; jessen 1918). from the end of the middle weichselain to the latest late weichselian, the area was subjected to terrestrial glaciation (houmarknielsen et al. 1996). after the ice melted back from the main stationary line, a marine environment was re-established and persisted until isostatic rebound resulted in subaerial exposure of the seabed of the younger yoldia sea (jessen 1918; figs 12, 13). a new lithostratigraphical subdivision is proposed to cover the three upper pleistocene successions (fig. 14). the systematic stratigraphic framework is based on formations defined according to the guidelines given by rawson et al. (2002). the lower marine unit, corresponding to the deposits representing the older yoldia sea and formerly referred to as the skærumhede series (jessen et al. 1910), is here defined as the skærumhede group (new group). the group includes the middle weichselian stortorn and lønstrup klint formations (new formations) and an unnamed lower unit mainly including the eemian and lower weichselian deposits (figs 15, 16). four formations are distinguished in the glacioterrestrial unit: the glaciofluvial and glaciolacustrine rubjerg knude formation (new formation), the kattegat till formation (houmark-nielsen 1987, 1999, 2003), the ribjerg formation (new formation) and the mid danish till formation (houmark-nielsen 1987, 1999, 2003). the uppermost major unit, comprising the post-glacial arctic marine younger yoldia clay and saxicava sand of jessen (1918, 1931), is referred to the vendsyssel formation (new formation) (fig. 14). skærumhede group new group history. the skærumhede group includes most of the lithological units formerly described as the skærumhede series (jessen et al. 1910). these include the marine eemian, the marine lower weichselian and the marine–brackish–lacustrine beds in the middle weichselian (figs 14, 15; lykke-andersen & knudsen 1991; knudsen 1994). recognition of the group is primarily based on a research borehole behind the farm at skærumhede, about 10 km west of frederikshavn (fig. 13), that was drilled by the geological survey of denmark to investigate the source of natural gas in the vicinity of frederikshavn (jessen et al. 1910). the well penetrated to a depth of 235 m and terminated in upper cretaceous chalk. above the chalk, a 20 m thick unit of till and glacial sediments was encountered. the till is now referred to the saalian (lykke-andersen 1987), and forms the basal unit of the quaternary succession over most of north jylland (fredericia 1982, 1983a, b; pedersen 1989). the succession above the saalian glacial sediments was described under the heading: ‘the marine skærumhede series’ by jessen et al. (1910 pp. 67, 156). this unit is c. 123 m thick, from 57.4 m to 180.3 m below surface, corresponding to a lower boundary at 157.1 m and a top at 34.2 m below sea level. it was subdivided into three biostratigraphic zones: 1) the turritella terebra zone (74 m thick), 2) the abra nitida zone (8.5 m thick) and 3) the portlandia arctica zone (40 m thick) (jessen et al. 1910). additional details were added to the unit based on several glaciotectonically dislocated outcrops in the northern part of vendsyssel by jessen et al. (1910) and jessen (1918, 1931). subsequent discussion concerning the stratigraphic position and dif ferentiation of the skærumhede series resulted in a new borehole, which was directed by the geological survey of denmark at the skærumhede locality in the early 1970s. although the borehole only went down to 120 m below surface, it gave a good record of the lithology and macrofauna and in particular provided samples for a detailed foraminiferal investigation (bahnson et al. 1974). facing page: fig. 14. schematic stratigraphic log of the units represented in the rubjerg knude glaciotectonic complex. the fossils indicated on the log represent 14c-dated samples. 34 name. the skærumhede group is named after the locality of skærumhede c. 10 km west of frederikshavn, denmark (fig. 13). type section. the type section is defined as the skærumhede well (dgu no. 10.4 and 10.392) (fig. 15), where the pioneer drill site for natural gas was situated at a barren and unfertile place caused by seepage of gas from the subsurface (fig. 13; jessen et al. 1910). reference sections. reference sections are proposed in well-documented borehole sections: the nørre lyngby ii well (dgu no. 8.137) described by lykke-andersen (1987), and the skagen iii well (dgu no. 1.287) recorded by knudsen (1994) and petersen (2004) (fig. 16). lithology. the skærumhede group consists of rather uniform bluish-black to dark grey clay with minor intercalations of silt and fine-grained sand. the silt laminae and thin fine-grained sand beds become more common towards the top of the group. macrofossils vendsyssel formation rubjerg knude formation lønstrup klint formation stortorn formation skærumhede group (undiff.) clay silty mud dropstones in mud sandy mud sand till gravel chalk sk æ ru m he de g ro up 10 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 210 220 230 clay silt sand cobble pebblef m c skærumhede well w ei ch se lia n u . c re ta ce ou s sa al ia n ee m ia n m below surface (e. 23 m a.s.l.) marine fossil fig. 15. lithostratigraphic log of the skærumhede well dgu no. 10.4, the type section of the skærumhede group and the stortorn formation. 35 are present through most of the group but decrease in abundance towards the top (jessen et al.1910). dropstones are present in the middle of the group and increase in abundance towards the uppermost part, in which graded silts and sands are intercalated with grey mud. boundaries. the lower boundary is the unconformity on top of the saalian till. the upper boundary is an unconformity overlain by coarse clastic sediments interpreted as a residual boulder bed (unit b of sadolin et al. 1997). thickness. the thickness of the group varies from nearly 130 m in the type section in the skærumhede well, to c. 48 m in the nørre lyngby well and c. 45 m in the skagen iii well (fig. 16). distribution. knowledge of the distribution of the group in the vicinity of rubjerg knude is based on the skærumhede well (dgu no. 10.4), the nørre lyngby well (dgu no. 8.137) and the skagen well (dgu no. 1.287; figs 13, 16). the group is also known from wells in the northern part of vendsyssel and the islands of læsø and anholt in the kattegat. according to these records, the group extends from the western part of the kattegat at frederikshavn and læsø, towards the south central part of vendsyssel, from where it continues offshore between anholt and djursland (knudsen 1994, fig. 3). the group extends offshore into the strait between læsø and the swedish coast. the southward extent is not known, but is probably up to about 30 km south of anholt. the extent to the north is also uncertain and has not yet been mapped. it is inferred that it may occur in the western part of the skagerrak (knudsen 1994) and it may also extend out into the northern part of the north sea. age. the age of the group extends from the beginning of the eemian, about 130 000 years b.p. (knudsen 1994), to the latest part of the middle weichselian, about 30 000 years b.p. (houmark-nielsen 1999). depositional environment. at the lower boundary of the group, red corroded flints were recognised in the skærumhede well (jessen et al.1910) indicating that the top surface of the saalian till had been exposed and subjected to subaerial erosion prior to the transgression that culminated in the eemian. during the eemian, a deep-water shelf environment was established with water depths exceeding 100 m; in the early weichselian, water depths decreased dramatically to less than 50 m (knudsen 1994). the decrease in water depth continued during the middle weichselian under increasing glacial influence. subdivisions. the upper skærumhede group is subdivided into the stortorn and lønstrup klint formations. the lower part of the group, mainly including the marine eemian and lower weichselian deposits, is presently undifferentiated. stortorn formation new formation history. in the lønstrup klint section, two units of grey-bluish clay subjected to glacial deformation have been distinguished, the diluvial clay and the portlandia arctica clay (jessen 1931). the latter unit corresponds to the so-called older yoldia clay (ældre yoldialer in danish), which in the skærumhede well was referred to as the portlandia arctica zone and in the crosssection of lønstrup klint is indicated to occur at three localities (jessen 1931). the most impressive of these is the stortorn site, where dark grey – black clay, rich in mollusc shells, crops out (fig. 21). the site is inaccessible, or difficult of access, since the slippery clays occur in the breaker zone at the foot of the almost vertical cliff section. the other two localities are the cliff sections just beyond the town lønstrup, locally named ‘lille blå’ (little blue), and the clif f section below the northern corner of the mårup churchyard. at all three sites, the unit is tectonically disturbed which hampers detailed logging of the succession. in addition, the formation occurs locally in the lower thrustsheet duplexes north of mårup church and in the moserende cliff section (see description of that section, below). name. the formation is named after the stortorn cliff section at lønstrup klint. the formation is here incorporated within the lowermost thrust unit in the stortorn section (see below). type section. the type section for the formation is the skærumhede well, dgu no. 10.4 (figs 14, 15). reference sections. the reference sections for the formation are the outcrops at stortorn and north of the northern corner of mårup churchyard (fig. 17) in the 36 weichselian saalian holocene upper cretaceous eemian vendsyssel fm kattegat till fm rubjerg kn. fm lønstrup kl. fm lønstrup kl. fm stortorn fm stortorn fm skærumhede group (undiff.) stortorn fm lønstrup klint fm clay silt sand cobble pebblef m c 10 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 10 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 210 220 230 clay silt sand cobble pebblef m c silt clay sand cobble pebblef m c nørre lyngby m below surface (e. 20 m a.s.l.) skærumhede m below surface (e. 23 m a.s.l.) skagen m below surface (e. 3 m a.s.l.) 10 20 30 40 50 60 70 80 90 100 110 120 130 140 150 v endsyssel fm clay silty mud dropstones in clay and silty mud sandy mud sand till gravel chalk marine fossils vend sys sel fm rubj erg kn. f m 37 mud sand f. m. c. gr.pb.co. 0 m 5 10 15 20 25 sandy mud sand dropstones in mud gravel shells lamination structureless climbing ripples trough cross-bedding ball-and-pillow / convolute bedding thrust fault clayey mud current ripple cross-lamination vendsyssel formation dune sand lønstrup klint formation sk æ ru m he de g ro up x aar 4069 stortorn formation unconformity thrust zone fig. 17. sedimentological log of the succession in the southern part of the ribjerg section (above the ‘store blå’). the stortorn formation records an arctic marine deposit, yielding shells typical of this environment: hiatella arctica, mya truncata and portlandia arctica. the boundary between the stortorn and lønstrup klint formations constitutes a thrust-fault breccia indicating dif ferentiation into thrust-fault duplex segments of the skærumhede group. the vendsyssel formation at the top of the section was deposited on an erosional unconformity with a lag conglomerate at the base. the location of the sample collected for 14c dating (aar 4069) is indicated (see table 2). facing page: fig. 16. simplified lithological logs from three thoroughly documented wells in vendsyssel (nørre lyngby: dgu no. 8.137; skærumhede: dgu no. 10.4; skagen: dgu no. 1.287, for location see fig. 13). the logs illustrate the stratigraphic correlation of the units defined in the investigation of the rubjerg knude glaciotectonic complex. the difference in thickness of eemian–weichselian deposits mainly reflects the average content of sand; the skagen well represents a deeper marine depositional environment compared to the skærumhede well. note that the cretaceous deposits in the skagen well comprise turonian–cenomanian greensands. the figure is based on information from jessen et al. (1910), bahnson et al. (1974), lykke-andersen (1987), lykke-andersen & knudsen (1991), knudsen (1994) and petersen (2004). 38 lønstrup klint coastal cliff. the nørre lyngby well (dgu no 8.137, lykke-andersen 1987) is the well closest to stortorn where the undisturbed formation has been penetrated. additional sections include the coastal clif f at hirtshals displaying allochthonous peats in the black clay formation, and the skagen iii well (dgu no 1.287; petersen 2004) that includes a clay unit, 8 m thick, here referred to the stortorn formation (fig. 16). lithology. the stortorn formation consists of black, locally dark grey – bluish structureless clay with a large number of dropstones, which are commonly glacially striated. lenses or irregular beds, up to 10 cm thick, of shell debris (gravel-size) occur scattered in the unit, and the abundance of shells in local patches gives the formation a white spotted appearance (fig. 18). at the top of the formation, the clayey mud changes colour from dark bluish grey to violet-brown and develops recognisable lamination. fossils. the unit has been referred to the portlandia arctica zone of jessen et al. (1910) since this is the most abundant mollusc species in the clay (bahnson et al. 1974). macoma calcarea is another common mollusc and hiatella arctica occurs in abundance. a list of characteristic molluscs and their distribution in the unit is given by k.s. petersen (in: bahnson et al. 1974). moreover, the presence of balanus sp. and additional erratic macrofossils are reported. the most common microfossils are the foraminifers elphidium excavatum and cassidulina crassa. p.b. konradi and k.l. knudsen (in: bahnson et al . 1974) documented and discussed the foraminiferal fauna. boundaries. the lower boundary of the formation is defined by a shift from clayey mud to mud with a marked increase in coarse-grained ice-rafted debris. the increased content of coarse-grained material is associated with an abundance of mollusc shells and fragments. the upper boundary of the formation is defined at the transition from marine clay showing diffuse lamination and colours varying from grey bluegreen to violet-brown, to a grey clayey and silty mud intercalated with graded silt and fine-grained sand laminae a few millimetres thick. thickness. the formation is about 20 m thick. in the skagen iii well, the formation is only about 8 m thick, probably due to the more offshore position and deeper water environment in this part of the basin (petersen 2004). distribution. the distribution of the stortorn formation is identical with the distribution of the skærumhede group. the formation can be readily identified in the nørre lyngby well (lykke-andersen 1987; figs 13, 16) and it has also been described from the hirtshals cliff section (lykke-andersen 1971). in addition, it is known from the deeper wells in the main part of the vendsyssel area and from the islands of læsø and anhiatella shells 5 cm fig. 18. the stortorn formation in the stortorn section is dominated by black sticky clay. locally, shells of hiatella arctica and portlandia arctica are very abundant. dropstones are also common in the formation. photograph: august 2001. 39 holt (fredericia 1982, 1983a, b, 1984; lykke-andersen 1987). age. three shell samples from the stortorn formation at lønstrup klint have been 14c dated, using the atomic mass spectrometric (ams) method, for the present investigation. two of the samples were derived from the archives of the former geological survey of denmark; two shells of hiatella arctica were chosen for dating the formation at the stortorn locality and from the northernmost outcrop of the formation at lønstrup klint (‘lille blå’, at the base of the northern part of the ribjerg section, collected and described by a. jessen) (table 2). the third sample was taken in 1996 and comprises shells of hiatella arctica from the shellbearing clay outcrop at stortorn (reference dgu no. 00136, aar-4069, table 2). these were dated to test the collection made nearly 100 years earlier and provided an age for the lowermost stortorn formation, namely 31 300 (± 400) years b.p. the age of the upper levels of the formation, as represented by the muddy sediments at the ‘lille blå’ section is slightly younger (30 000 (± 400) years b.p. (table 2)). the new dating of the stortorn formation corresponds well with previous age dates from the upper part of the skærumhede group, which gave 32 000 years b.p. (seidenkrantz & knudsen 1993). depositional environment. the presence of a boreal fauna including mytilus edulis, arctica islandica and zirphaea crispata in the shell-debris gravel in an environment characterised by a bottom fauna of portlandia arctica and macoma calcaria led nordmann (1928) and jessen (1931) to conclude that the boreal shallow-water faunas of interglacial affinity were transported as ice-rafted material into more offshore arctic marine environments. the most convincing examples of such erratic material are the dropstones with balanus sp. the stortorn formation is thus interpreted to have been deposited during a period of decreasing water depths in a marine environment characterised by dispersal of erratics from drifting icebergs. lønstrup klint formation new formation history. the characteristic development of this formation, viz. grey clayey muds interbedded with layers of fine-grained sand, occurs in the steeply inclined sheets that are prominent in the cliff of lønstrup klint. due to the absence of macrofossils, this unit was named diluvialler (diluvial clay) and the unit was correlated with the uppermost part of the skærumhede group in the skærumhede well (jessen et al. 1910; jessen 1918, 1931). the sedimentology of the formation was described from sandrende at lønstrup klint under the heading unit a by sadolin et al. (1997). name. the formation is named after the coastal cliff of lønstrup klint. type section. the type section for the formation is at sandrende in lønstrup klint, situated between point 3500 and 3600 m in the rubjerg knude cross-section (plate 1), from where a sedimentological log was provided by sadolin et al. (1997) (figs 13, 19). reference sections. reference sections are defined at ulstrup rende (fig. 20) and at kramrende (fig. 21) situated at 5950 m and 4500 m, respectively, in the rubjerg knude cross-section (plate 1). moreover, the skærumhede and the skagen wells are reference sections for the western and northern development of the formation (fig. 16). lithology. the formation consists of blue-grey clayey and silty laminated mud and cross-laminated beds of fine sand. the lowest part of the formation is characterised by dark grey mud, interlayered with laminated to thin-bedded clayey and silty mud. the light grey beds, 1–5 cm thick, grade upwards from light grey silt to dark grey clay (fig. 22). some of the dark grey clayey mud levels are interbedded with thin lenticular, light-coloured silt and fine-grained sand laminae (fig. 23). silty to fine-grained sandy beds may be up to 1 m thick. dropstones occur scattered in the bluegrey mud. in the upper part of the formation, beds of light grey sand, 3–8 m thick, occur interbedded with a few thin beds of laminated mud. the thick sand beds are characterised by climbing ripple cross-lamination. in the cliff section at rubjerg knude, much of the primary bedding in the lønstrup klint formation is disturbed by water-escape structures (ball-and-pillow etc.) and hydrodynamic brecciation (flame to diapir structures). fossils. macrofossils have not been found in the formation and the foraminifers, dominantly elphidium excavatatum, are interpreted to be redeposited (lykkeandersen 1987). 40 mud sand f. m. c. gr.pb.co. 5 10 15 0 m l/r-unconformity 20 25 30 35 40 sandy mud sand gravel basal till lamination structureless climbing ripples trough cross-bedding slump structures ball-and-pillow / convolute bedding thrust fault rubjerg knude formation lønstrup klint formation kattegat till formation clayey mud current ripple cross-lamination fig. 19. sedimentological log of the succession in the sandrende section, including the type sections of the lønstrup klint and rubjerg knude formations. the lønstrup klint formation represents lacustrine deposition, whereas the rubjerg knude formation records a shift from fluvial to lacustrine sedimentation, returning to fluvial sedimentation in the upper levels (sadolin et al. 1997). note the synsedimentary small-scale thrust structures that appear in the upper levels of the rubjerg knude formation indicating that the formation was deposited in a piggyback basin. 41 boundaries. the lower boundary of the lønstrup klint formation is placed at the transition from the typical marine clay with a macrofossil fauna of the stortorn formation to unfossiliferous laminated clayey and silty muds with laminae and thin beds of fine sand. the upper boundary is the marked erosional unconformity between the lønstrup klint and rubjerg knude formations, referred to as the l/r-unconformity (figs 14, 24). thickness. the maximum thickness of the formation is about 25 m, but with large variations due to erosional relief at the l/r-unconformity. distribution. the lønstrup klint formation is distributed over the main part of vendsyssel. it is erosionally truncated at the top towards the south where underlying glacial deposits or upper cretaceous chalk constitute the surface geology of the coastal areas north of limfjorden. to the east, it probably extends offshore into the middle part of kattegat from where it is known in wells on the islands of læsø and anholt (lykke-andersen 1987; knudsen 1994). the extent out into the north sea to the west remains unknown. towards the north, it extends offshore into the skagerrak beyond skagen, where it is recorded in the skagen iii well (petersen 2004). 24 20 15 10 5 0 m mud sand f. m. c. gr. pb.co. glacitectonite rubjerg knude formation ulstrup glaciofluvial beds lønstrup klint formation clayey mud sandy mud sand gravel lamination current ripple cross-lamination trough cross-bedding tectonite water-escape pipes and sand-filled cracks structureless sediment l/r-unconformity fig. 20. sedimentological log of the succession in the southernmost thrust sheet in the ulstrup section. the log was measured at ulstrup rende, situated at point 5980 m in the cross-section (plate 1). note the 4 m thick sand-crack dominated thrust zone that characterises the lower part of the lønstrup klint formation, which developed during thrust-sheet translation along the hangingwall flat (figs 46, 47). note also that in the rubjerg knude formation, the lowermost 7 m corresponds to the glaciofluvial ‘ulstrup beds’ (figs 28, 29). 42 age. 14c dating of plant debris from the formation indicates an age of about 30 000 b.p. (houmark-nielsen et al. 1996). this age is compatible with the 32 000 b.p. age derived from the underlying stortorn formation. depositional environment. the transition from the stortorn formation to the non-fossiliferous lønstrup klint formation is interpreted as a shift from a normal, arctic marine environment through brackish to a freshwater environment dominated by rapid deposition of suspended sediment supplied to the basin by meltwater. the sharp-based normally graded silt and sand beds are interpreted as fine-grained turbidites. the dark mud was deposited from suspension, whereas the lenticular sand/silt laminae represent wave-reworked, distal storm-sand layers. sedimentation started below storm-wave base and it is suggested that the lake environment was deep and of fairly wide extent. the occurrence of numerous fine-grained sandy turbidites sourced mainly from the south probably reflects exposed land areas in the southern part of vendsyssel during the low stand of sea level (sadolin et al. 1997). the sand beds are interpreted to record relatively rapid sedimentation by sediment gravity 0 m 5 10 15 20 25 thrust zone rubjerg knude formation lønstrup klint formation sandy mud sand gravel lamination structureless climbing ripples trough cross-bedding ball-and-pillow / convolute bedding tectonite clayey mud current ripple cross-lamination mud sand f. m. c. gr.pb. co. l/r-unconformity fig. 21. sedimentological log of the lønstrup klint and rubjerg knude formations in the kr01 thrust sheet in the southern part of the kramrende section (point 4500 in plate 1). the fine-grained, thinto medium-bedded sandy turbidites in the lønstrup klint formation are interbedded with thin layers of blue-grey silty mud. these sand beds are often disrupted into ball-and-pillow load structures (see fig. 48). note the tectonite at the base of the succession, related to the hanging-wall flat of the kr01 thrust sheet. 43 flows in a glaciolacustrine environment; the thickest of these beds may represent deposition within one summer of sediment derived from the southern slopes of the basin (sadolin et al.1997). hydrodynamic deformation of the strata was initiated at a very early stage in the glaciotectonic process, as loading by the superposed rubjerg knude formation and by glaciotectonic thrust sheets resulted in increasing pore-water pressure. hydrodynamic brecciation continued during deformation until the displacements of thrust sheets ceased. upper weichselian lithostratigraphic units rubjerg knude formation new formation history. in the steep cliff section at rubjerg knude, the thrust sheets, consisting of the grey-blue coloured clay of the lønstrup klint formation, are depositionally overlain and structurally underlain by light-coloured yellowish sand, named diluvialsand (diluvial sand) by jessen (1918, 1931). the succession was refig. 22. laminated to thin-bedded clayey and silty mud in the lower part of the lønstrup klint formation in the rubjerg knude fyr section. the bedding is defined by layers grading from light grey silt to dark grey clay, and the sharp-based normally graded beds are interpreted as fine-grained turbidites. the coin for scale is 2.5 cm in diameter. photograph: september 1985. fig. 23. dark grey clayey mud interbedded with thin lenticular, light coloured silt and fine-grained sand laminae. the mud was deposited from suspension, whereas the lenticular laminae represent wave-reworked, distal storm-sand layers, deposited below storm-wave base. photograph: june 1993. 44 fig. 24. the l/r-unconformity is steeply inclined in the gr08 thrust sheet. note the large-scale cross-bedding in the basal unit of the rubjerg knude formation, which onlaps the unconformity (r-onlap). photograph: june 1993; rucksack for scale. l/r-unconformity 0 m 13 10 ulstrup glaciolacustrine beds lønstrup klint formation rubjerg knude formation thrust zone 5 mud sand f. m. c. gr.pb. co. sandy mud sand gravel lamination structureless sediment trough cross-bedding ball-and-pillow convolute bedding water-escape pipes and sand-filled cracks thrust fault clayey mud current ripple cross-lamination fig. 25. sedimentological log of the succession in the northern thrust sheet in the ulstrup section. the log was measured near the ulstrup steps at point 5625 m in the cross-section (plate 1). the base of the log is the hanging-wall flat of the ul02 thrust sheet and the lowermost 2 m constitute the thrust zone. the boundary between the lønstrup klint formation and the ulstrup glaciolacustrine beds above (lower unit of the rubjerg knude formation) is a flat, non-erosional surface but can be traced to the northern part of the ul02 thrust sheet where this boundary is a clear erosional unconformity. 45 ferred to as units b–d in the sedimentological study by sadolin et al. (1997). name. the formation is named after rubjerg knude, the highest part of the lønstrup klint cliff.. type section . the type section is at sandrende in the lønstrup klint cliff section (figs 13, 19). reference section. four reference sections are defined, all situated in the vicinity of rubjerg knude. in the distal part of the rubjerg knude glaciotectonic complex, two reference sections are defined at ulstrup. the first of these is located in ulstrup rende at point 5900 m in the rubjerg knude cross-section (plate 1) which demonstrates the presence of coarse-grained glaciofluvial channel fill deposits (fig. 19). the second reference section at ulstrup is located at point 5450 m in plate 1 and documents the occurrence of fine-grained clayey muddy glaciolacustrine beds in the formation (fig. 25). the third reference section is situated at martørv bakker (point 4850 m in plate 1), which demonstrates diamictitic sediments including slump units in a piggyback basin (fig. 26). the fourth reference section is located at moserende at point 1750 m in plate 1. this section illustrates the formation in a piggyback basin situated in a proximal position in the rubjerg knude glaciotectonic complex (fig. 27). 23 mud sand f. m. c. gr.pb. co. 20 15 10 5 0 m vendsyssel formation holocene peat ( martørv ) aeolian sand rubjerg knude formation (including diamict sediments and slump-folded units) lønstrup klint formation l/r-unconformity erosional unconformity sandy mud sand gravel lamination structureless sediment trough cross-bedding mud with scattered pebbles and cobbles slump structures ball-and-pillow convolute bedding clayey mud current ripple cross-lamination fig. 26. geological log of the diamict sediments and slump-fold structures, which represent the rubjerg knude formation in the piggyback basin on the back of the mb02 thrust sheet in the martørv bakker section, point 4880 in the cross-section (plate 1). 46 lithology. the dominant lithology of the formation is fineto medium-grained sand. beds of gravel occur in the lowermost 1–5 m, related to the initial deposition succeeding the formation of the erosional unconformity (the l/r-unconformity) (figs 24, 27, 28, 29). the sediment source was partly the main central part of the danish basin, indicated by the content of 23–25% flint and upper cretaceous chalk, and partly outwash material from the propagating ice margin, indicated by the c. 75% basement clasts (jessen 1931). many sand beds display small-scale current ripple lamination, and some show well-developed climbing ripples (sadolin et al. 1997). large-scale cross-bedded sand is observed in shallow channel fills, and some trough cross-stratification occurs in relation to growth-fault structures formed along normal faults or depressions related to the formation of synsedimentary footwall synclines (see fig. 87). a series of large-scale accretionary cross-stratification structures are related to a shift in the substratum inclination during thrust-fault propagation (see figs 82, 95). clasts of clay derived from the lønstrup klint formation are common, and in some of the syntectonic settings these beds rich in clay-clasts may be regarded as sedimentary clastic breccias with olistoliths or lumps of sandy mud 1–5 m in size (fig. 56). the olistoliths represent the frontal parts of thrust sheets, which gravity-glided out into depressions formed during thrust-fault propagation. locally, some of the depressions developed into small glaciolacustrine basins characterised by interbedded finegrained sands and sandy muds with current crosslamination; these deposits may reach a thickness of up to 5 m (fig. 25). fossils. redeposited fossils occur together with accumulations of twigs and amber (‘ravpindelag’). wellpreserved arctic mosses suitable for 14c dating the formation have been separated from the organic debris (houmark-nielsen et al 1996). in the basin at stensnæs, a large number of the mollusc shells were rel/r-unconformity mud sand f. m. c. gr.pb. co. 0 m 5 10 15 20 sandy mud sand gravel lamination structureless sediment climbing ripples trough cross-bedding slump structures ball-and-pillow convolute bedding anastomosing joints thrust fault rubjerg knude formation lønstrup klint formation clayey mud current ripple cross-lamination fig. 27. geological log of sediments and thrust faults in the mr03 thrust sheet in the moserende section (point 1750 in plate 1). 47 garded as a redeposited interglacial fauna by jessen (1931). among the shells are astarte sp., cardium sp., arctica islandica, leda pernula, mya truncata, hiatella arctica; a full list of this diverse fauna is given in jessen (1931, p. 63). an astarte sp. shell (aar-4066) was 14c dated to 43 000 ± 1300 years b.p. (table 2), which must be regarded as close to an infinite age, thus supporting the suggestion of jessen (1931) that these shells represent redeposited interglacial faunas. boundaries. the lower boundary of the formation is placed at the erosional l/r-unconformity capping the lønstrup klint formation. this has a relief of 0.5–1 m and is commonly overlain by an up to 0.5 m thick clast-supported residual gravel bed. in the distal southern part of the rubjerg knude glaciotectonic complex, the l/r-unconformity is located close to sea level, and is a convenient structural reference level. it represents the top level of pre-tectonic sedimentation, and is hence also a reference surface for the construction of the balanced cross-section. the upper boundary is placed at the glaciotectonic unconformity below the kattegat till formation. thickness. the thickness of the formation is about 25 m, but it varies considerably according to local depositional and erosional development. fig. 28. the glaciofluvial ulstrup beds deposited above the l/r-unconformity (l/r-u) on top of the lønstrup klint formation. note the boulder in the lowermost part of the glaciofluvial ulstrup beds indicating the high-energy (upper flow regime) of the meltwater streams that deposited the beds (compare with fig. 20). the divisions on the measuring pole are 20 cm. photograph: may 1998. ulstrup section, 5950 m in cross-section (see plate 1). fig. 29. the glaciofluvial ulstrup beds with ‘fossil frozen’ sand clasts that indicate ground-frozen conditions in the source area of the sand clasts; they were probably derived from the lower part of the rubjerg knude formation farther north. photograph: may 1998. 48 distribution. the rubjerg knude formation was mainly deposited and preserved between the thrust sheets of the rubjerg knude glaciotectonic complex. the formation extends towards the south to the area around nørre lyngby where it was mapped as ‘morænesand’ (moraine sand – sandy till) by jessen (1918, 1931), and it has not been identified south of løkken. the formation is not recognised in the area north of lønstrup, which was mainly covered by ice during deposition of the formation. to the east it can be traced in wells about 10 km inland, where it pinches out due to erosion during the transgression of the younger yoldia sea. the formation probably does not extend out into the north sea to the west since it is largely situated above sea level. age. the formation has an age range of 30 000 – 20 000 years b.p. based on 14c dating of mosses, separated from the organic debris draping the ripple lamination, and twigs and amber layers (houmark-nielsen et al. 1996; table 2). the mosses investigated were transported from a carbonate-rich source area, probably the cretaceous chalk outcrops near limfjorden (fig. 13). the time span for redepositing plant debris is not regarded to exceed hundreds of years, and the age of the formation was thus interpreted to be closer at 29 000 than 30 000 years b.p. (fig. 13; houmark-nielsen et al. 1996). depositional environment. the rubjerg knude formation is interpreted to have been deposited on an outwash plain, which was dissected into smaller piggyback basins during glaciotectonic thrust faulting. during the development of the piggyback basins, deposition was controlled by the propagation of the thrust sheets. north of lønstrup, a large depression is regarded as the hole in a hill-and-hole pair from where the piggyback basins contemporaneous with deposition of the rubjerg knude formation were dislocated to the south during the glaciotectonic deformation. kattegat till formation history. the rubjerg knude formation is truncated by a glaciotectonic unconformity and overlain by the kattegat till formation (fig. 14). the formation was erected by houmark-nielsen (1987) in the areas surrounding the southern part of the kattegat and is interpreted to have been deposited during the weichselian ice advance from norway. subsequent studies have demonstrated that the formation can be identified over much of the northern part of the danish basin (fig. 12; houmark-nielsen 1999, 2003). name. the formation is named after the kattegat strait (fig. 12). type section. the type section is at hundested klint (fig. 12; houmark-nielsen 1987). reference section. two reference sections are defined in the lønstrup klint coastal section, namely the top of the sandrende locality at point 3700 m in plate 1 (figs 19, 30) and the cliff exposure c. 400 m north of the mårup church, at point 500 m in plate 1 (fig. 31). lithology. at the type section, the formation is a grey, clayey till only a few metres thick; the erratic clasts are dominantly crystalline rocks of fennoscandian provenance and palaeozoic limestone. foraminifers and shell fragments in the matrix have been identified as having been derived from the skærumhede group (houmark-nielsen 1987). in the rubjerg knude area, the till is light beige-brown weathering, dark grey and sandy with fineto medium-grained sand in the matrix. erratic pebbles and cobbles occur scattered in the matrix, and indicator pebbles of permian porphyry from the oslo region are common (1–5% of the erratics). in the main part of rubjerg knude cliff section, the formation drapes the glaciotectonic complex; over large areas, it has been subjected to aeolian erosion that has removed the fine-grained matrix and left the erratics as a cobble pavement. north of the mårup church, the formation comprises a shear till with erratics interlayered in a glaciotectonic breccia dominated by shear-deformed clayey mud derived from the top of the skærumhede group (fig. 31). in the northern part of the lønstrup klint cliff section (the ribjerg section), the kattegat till formation is absent. the glacial advance, represented elsewhere by the kattegat till formation, is here recorded only by a glaciotectonic unconformity and an underlying glacitectonite characterised by a dense anastomosing framework of joints penetrating the skærumhede group (fig. 32). boundaries. the lower boundary of the formation is the glaciotectonic unconformity formed by the shear at the base of the advancing norwegian ice. below the unconformity, a glacitectonite 1–2 m thick developed due to shear deformation of the clay and sand in the lønstrup klint and rubjerg knude formations. 49 the upper boundary is the subaerial erosional surface above the 1.5 m thick sandy till, commonly reduced to a 0.25 m thick residual pavement. thickness. the formation is up to 1.5 m thick at rubjerg knude. distribution. the formation has been recognised from the central and northern part of the west coast of jylland and vendsyssel over djursland and sjælland to hven and glumslöv in the western part of skåne, sweden (fig. 12; houmark-nielsen 2003). age. the age of the kattegat till formation is bracketed by the lønstrup klint formation beneath (29 000 years b.p.) and the ribjerg formation above (26 000 years b.p.) (fig. 14, tables 2, 3). the age is estimated to be 27 500 ± 1000 years b.p. (houmark-nielsen 2003). depositional environment. the kattegat till formation is interpreted as a lodgement till. in the area between lønstrup and mårup church, the diamict lithology of the upper skærumhede group suggests that deformation of the substratum below the glaciotectonic unconformity was initiated by mud-mobilisation of water-saturated clay, silt and fine-grained sand. material, including erratic clasts from the lodgement bed along the sole of the ice, dropped into the mudmobilised unit. during the advance of the ice, the mudmobilised zone became consolidated, and sub-horizontal anastomosing joints formed in the substratum. the depositional environment therefore changed from a wet-based glacial advance to an advance over dehydrated or even frozen substratum during the deposition of the kattegat till formation at rubjerg knude. fig. 30. the c. 1 m thick sandy till on top of the sandrende section is referred to the kattegat till formation. the maximum size of the erratic clasts is 25 cm. the marked planar erosion sur face above the till was initially formed by glacial truncation, which subsequently was exposed to aeolian erosion and finally covered by dunes. photograph: july 1993. 50 ribjerg formation new formation history. the ribjerg formation is a new formation proposed for the c. 25 m thick glaciofluvial sand unit that crops out between the northern part of the lønstrup klint clif f section and the northern part of the town of lønstrup. the unit was indicated in the northernmost c. 2 km of the cross-section of jessen (1931), but it was only regarded as part of the main diluvialsand (diluvial sand). since the formation was deposited in the late weichselian between ice advances from norway and central sweden, it might in a glaciodynamic context be correlated with the outwash deposits of the tebbestrup formation in djursland (larsen et al. 1977; pedersen & petersen 1997). name. the formation is named after the hill of ribjerg at lønstrup. type section. the type section is located at the cliff below the ribjerg hill, south-west of lønstrup (figs 13, 33). lithology. the ribjerg formation is characterised by fineto medium-grained sand showing large-scale trough and channel cross-stratification (fig. 33). at fig. 31. a pocket of sandy till overlying a glacitectonite and associated features of subglacial deformation. the till is referred to the kattegat till formation. the locality is situated c. 450 m north of mårup kirke. photograph: july 1994. fig. 32. planar-parallel and elongated anastomosing shear joints that are typical of the glacitectonite at the top of the skærumhede group below the blåunconformity in the ribjerg section. photograph: july 1994. 51 the base of the formation, tabular grey mud clasts (1 × 5–10 cm in cross-section) form a lag deposit in the fine-grained sand. lamination outlined by heavy minerals occurs in the lowermost metre of the formation and current-ripple lamination with mud chips is also present (fig. 33). the lower part of the formation is characterised by horizontal planar laminated finegrained sand interlayered with thin beds (0.1–0.5 cm thick) of clay-draped current ripples. this facies is overlain by 0.5 m thick beds of fineto coarse-grained trough cross-bedded sand. troughs or channels, 2–5 m deep and 10–15 m wide, occur in the middle and upper part of the formation (fig. 111). in the central part of the troughs, the fill shows large-scale crossstratification. towards the margin of the troughs, the beds decrease in thickness and display small-scale mud sand f. m. c. gr.pb. co. 25 20 15 10 5 0 m ribjerg formation peat and dune sand mid danish till formation glaciotectonic-unconformity blå-unconformity skærumhede group x r-990222 x r-990223 x r-990224 sandy mud sand dropstones in mud gravel basal till shells lamination structureless mud clasts climbing ripples trough cross-bedding slump structures ball-and-pillow / convolute bedding water-escape pipes and sand-filled cracks tectonite clayey mud current ripple cross-lamination x aar 4067 fig. 33. sedimentological log of the ribjerg formation (type section) in the northernmost part of the ribjerg section. the formation was deposited above the blå-unconformity on top of the skærumhede group, and it is overlain by the sandy mid danish till formation. the ribjerg formation represents glaciofluvial deposition related to a channel-eroded foreland of an advancing glacier. note the sand-dykes that are interpreted to have formed by discharge of pore water due to high stream velocity. samples collected for optically stimulated luminescence dating are indicated, together with their laboratory numbers (see table 3). 52 current lamination with ripples draped by organic debris. the steeply inclined slopes (up to 30°) of the trough margins strike 88–94°, indicating an east–west current direction. erosional surfaces with slumped beds and pockets of gravel recur every 1 to 3 m (fig. 33). one of the most characteristic features of the formation is the large number of sand dykes and waterescape pillars, which are 5–15 cm wide and can be traced vertically for more than 1.5 m (fig. 34). the uppermost 3 m of the formation comprises thick gravel beds just below the flow till related to the mid danish till formation. boundaries. the lower boundary of the formation is defined at the erosional unconformity forming the top of the kattegat till formation or, where the erosion level penetrates deeper, the lønstrup klint formation. in the northern part of the lønstrup klint section, the unconformity on top of the ‘lille blå’ forms the lower boundary. the upper boundary is placed at the base of the flow till that forms the lower part of the mid danish till formation (figs 14, 33). thickness. the formation is about 25 m thick. distribution. the formation is only recognised in the fig. 34. sand dyke intruded in the glaciofluvial succession of the ribjerg formation. the sand dykes are interpreted to have formed by pore-water discharge from the sediment due to high velocity current flux through the channels. photograph: july 1994. ribjerg formation, type section. 53 vicinity of lønstrup and towards vennebjerg to the east. it is inferred to have been deposited over a larger area of north-western vendsyssel, which is now covered by the vendsyssel formation (see below). age. three samples were collected from the lower, middle and upper part of the formation for optically stimulated luminescence dating (r-990222, r-990223, r-990224; table 3); these samples indicate an age of 26 000 – 25 000 years b.p. depositional environment. the formation was deposited in fluvial channels cut by westward-flowing meltwater. the sand dykes and water-escape pillars are indicative of high pore-water pressure due to rapid deposition and very fast meltwater flux through the channels. the outwash deposits are interpreted as a valley sandur that formed in the depression resulting from the hole left in the hinterland of the rubjerg knude glaciotectonic complex. the source of the meltwater was the ice margin of the advancing ice from central sweden in the late weichselian. mid danish till formation history. the ribjerg formation of the lønstrup klint section is overlain by a 3 m thick grey-brown till that is referred to the mid danish till formation (houmark-nielsen 1987, 1999, 2003). the formation was erected by houmark-nielsen (1987) to encompass tills deposited in the southern and central part of denmark during the weichselian ice advance from central sweden. it is known from the main part of the danish basin east and north of the main stationary line (fig. 12), and the records of its distribution in the northern part of denmark have recently been summarised by houmark-nielsen (1999, 2003) (figs 1, 12). name. the name of the formation reflects the prominent nature of this surface deposit in central (mid) denmark (fig. 12). type section. the type section is at ristinge klint on the island of langeland (houmark-nielsen 1987). reference section. a reference section is herein defined at ribjerg, sw of lønstrup (figs 13, 33). lithology. at the type locality, the mid danish till formation is a 5–8 m thick unit with at least two boulder pavements displaying ne–sw orientated glacial striation (sjørring et al. 1982). the formation is a grey to brown mostly clayey massive till with about 50% crystalline erratics of fennoscandian provenance; indicator clasts from the central eastern part of sweden (kinne-diabase) and from jurassic sedimentary rocks situated offshore in the kattegat (pedersen & petersen 1997) are abundant (houmark-nielsen 1987). in the reference section at ribjerg, the formation is table 3. optically stimulated luminescence dates on quartz, rubjerg knude and nørre lyngby, vendsyssel, northern denmark stratigraphic unit locality lab. id no. material age ka b.p. dose+ (gy) w.c.‡ (%) ref.* vendsyssel fm nørre lyngby r-829202a marine clay 16 ± 1 38.9 ± 1.3 30 (1) vendsyssel fm nørre lyngby r-829203 marine clay 17 ± 2 46.0 ± 0.8 29 (1) ribjerg fm ribjerg r-990224 fluvial sand 25 ± 2 45.3 ± 0.8 25 (3) ribjerg fm ribjerg r-990223 fluvial sand 26 ± 1 52.6 ± 1.5 26 (3) ribjerg fm ribjerg r-990222 fluvial sand 26 ± 1 53.1 ± 1.2 20 (3) lønstrup klint fm sandrende r-970204 fluvial sand 29 ± 2 57.6 ± 1.8 21 (2) stortorn fm ribjerg r-970203 lacust. sand 30 ± 2 65.0 ± 1.8 25 (2) + equivalent gamma dose. ‡ water content (saturation). * references: 1: strickertson & murray (1999); 2: houmark-nielsen (2003); 3: this study. 54 a yellow-brown weathering, grey-brown, sandy till comprising a lower stratified unit and an upper massive unit (fig. 35). the lower unit comprises laminated to finely bedded, matrix-supported diamictite with scattered pebbles. the matrix is fine-grained sand and the lamination and bedding are slump folded with n– s-trending fold axes and e-dipping axial planes, indicating a westward flow direction. the upper unit is a massive, structureless matrix-supported diamictite. erratic pebbles and cobbles are abundant and the till has a pronounced a-axis clast fabric dipping at low angles (c. 3°) towards the east (100°), indicating a shear transport direction towards the west. boundaries. in the reference section, the lower boundary of the formation is placed at the depositional conformity on top of the ribjerg formation where planar horizontal gravel beds are overlain by slumpfolded diamictites dominated by debris flow layers. the upper boundary is an erosional unconformity separating the diamictites from silt-streaked muds at the base of the vendsyssel formation. thickness. the formation reaches a thickness of more than 10 m at the type section at ristinge klint, but it is only 3 m thick at ribjerg in the reference section. distribution. in the rubjerg knude area south of rifig. 35. the sandy till that overlies the ribjerg formation is divided into a lower flow till and an upper lodgement till. the flow till is characterised by slump-folded debris flow lamination indicating flow from east to west. the till unit is referred to the mid danish till formation, which was deposited by the ice advance from the east, probably about 24 000 b.p. photograph: may 1985; notebook for scale. ribjerg formation, type section. 55 bjerg, where the ribjerg formation is absent, it has not been possible to differentiate occurrences of the mid danish till formation from the older kattegat till formation. however, the distribution is well documented throughout the danish basin east and north of the main stationary line (figs 1, 12; houmark-nielsen 1999, 2003). age. the age of the mid danish till formation is bracketed by the age of the ribjerg formation beneath (26 000 years b.p.) and the vendsyssel formation above (16 000 years b.p.) (fig. 14, tables 2, 3). the age is estimated to be 24 000 – 20 000 years b.p. (houmark-nielsen 2003). depositional environment. the lower unit is interpreted as a flow till deposited as debris flows from an ice margin to the east, prior to the ice advance towards the west. the upper unit is interpreted as a lodgement till deposited at the sole of the ice during the ice advance from central sweden towards the main stationary line situated in the central part of the north sea (fig.12). vendsyssel formation new formation history. north and south of rubjerg knude, the mid danish till formation is overlain by a succession of glaciomarine heteroliths, which are here defined as the vendsyssel formation. these deposits were mapped by jessen (1899), who related them to deposition in the younger yoldia sea in vendsyssel. jessen (1918) regarded the various facies of the vendsyssel formation as four stratigraphic units named the lower saxicava sand, the yoldia clay (usually prefaced ‘youngaeolian sand vendsyssel formation lønstrup klint formation glacitectonite 0 m 5 10 15 sandsilt m.f.si. c.clay gr. co. tectonite climbing ripples current ripple cross-lamination sand wave ripple lamination lamination trough cross-bedding gravel sandy mud clayey mud storm sand bed bioturbation shells imbricated mud clasts ball-and-pillow convolute bedding fig. 36. sedimentological log of the vendsyssel formation at the type section at stensnæs. the section is located at point 5120 in the cross-section (plate 1); the base of the log is 15 m a.s.l. 56 er’ to distinguish it from the older yoldia clay), the upper saxicava sand and the zirphaea beds. name. the formation is named after the region of vendsyssel in north denmark (figs 12, 13). type section. the type section is the coastal cliff section at stensnæs c. 1 km north of nørre lyngby in the central part of the west coast in vendsyssel (figs 13, 36). reference sections. two coastal cliff sections, north and south of rubjerg knude, are defined as reference sections. the locality to the north is the coastal clif f c. 500 m south of lønstrup, where heteroliths characterised by hiatella burrows crop out (fig. 37). to the south, the coastal cliff at nørre lyngby (north and south of the ramp leading down to the beach) probably gives the thickest accessible outcrop of the formation (fig. 38). this locality is furthermore close to the reference well dgu no. 8.137, where the maximum thickness of the formation is recorded (lykkeandersen 1987). lithology. two main lithologies dominate the formation: dark bluish-grey, clayey mud in the lower part and yellowish weathering light grey stratified heteroliths in the upper part. at the base of the formation, coarse-grained sands and gravels overlie the erosional unconformity above the mid danish till formation or older deposits (figs 36–38). the unit referred to as the lower saxicava sand by jessen (1918) is less than 2 m thick and is only present locally. accumulations of shell debris occur in places. in general, marine clayey mud forms the lower c. 6 m of the formation resting tectonite basal till aeolian sand vendsyssel formation kattegat till formation lønstrup klint formation glacitectonite 0 m 5 10 15 sandy mud sand gravel lamination wave ripple lamination storm sand bed bioturbation shells trough cross-bedding slump structure clayey mud sand m.f.si. c.clay gr.pb. co. fig. 37. sedimentological log of the vendsyssel formation at the reference section, situated halfway between lønstrup and mårup church at point 500 in the cross-section (plate 1); base of the log is 10 m a.s.l. the slump structure recognised in the lower part of the section (at about 4 m) is interpreted to have been produced by a grounding iceberg. the abundant shells in the section are hiatella arctica and the bioturbation was due to the infaunal activity of these molluscs (see fig. 41). 57 directly on the lower erosional boundary (figs 36– 38); dropstones, locally to boulder size, occur in the lower part of the clayey mud unit (figs 36, 37). the unit is highly impermeable such that groundwater wells out at the top of the clayey mud outcrops, often obscuring the exposures of the basal lithologies. above the clayey mud unit, horizontal stratified heteroliths form a unit 6–12 m thick. in places, the heteroliths grade into sandy mud characterised by wave ripple lamination (fig. 40). dark grey laminated mud is interbedded with fine-grained sand beds up to 10 cm thick in which wave ripple lamination is common. at the reference section, south of lønstrup, the heteroliths are intensively bioturbated by vertical trace fossils produced by hiatella arctica and the shells are often preserved in life position (fig. 41). the reference section at nørre lyngby is located in a half-graben structure with the steepest normal fault (dipping c. 60°s) situated north of the village (lykkeandersen 1992). south of nørre lyngby, the erosional unconformity below the vendsyssel formation dips 5–8° to the north. the beds above the unconformity are characterised by sedimentary breccias of mud clasts probably derived from the lønstrup klint formation (fig. 38). in the middle part of the formation, the beds are displaced by synsedimentary faulting (fig. 38) indicating that the half-graben formed during the deposition of the vendsyssel formation. aeolian sand vendsyssel formation kattegat till formation lønstrup klint formation sandsilt m.f.si. c.clay gr. co. 0 m 5 10 15 20 dropstones in mud tectonite climbing ripples current ripple cross-lamination basal till sandy mud sand lamination trough cross-bedding clayey mud imbricated mud clasts ball-and-pillow / convolute bedding normal fault thrust fault fig. 38. sedimentological log of the vendsyssel formation at the reference section, situated c. 350 m south of nørre lyngby. the beds rich in imbricated mud clasts reflect the tectonically active nature of the half-graben in which the section is located; the synsedimentary tectonic activity is further documented by the intraformational normal and thrust faults that occur in the lower half of the formation at nørre lyngby. 58 in the area north-east of hirtshals, the clayey mud is overlain by coastal sands, the so-called zirphaea beds (jessen 1918), and in the eastern part of vendsyssel the uppermost part of the formation comprises coarsegrained sand and gravel deposited in a large spit system (nielsen et al. 1988). fossils. the fossils characteristic of the formation are the molluscs portlandia arctica, hiatella arctica and zirphaea crispata. boundaries. the lower boundary is the erosional unconformity on top of the mid danish till formation or older deposits. the upper boundary is the top surface of the landscape upon which locally lie terrestrial deposits such as the allerød peat beds in the nørre lyngby bog (jessen & nordmann 1915), the boreal peat at martørv bakker (jessen 1931) and recent aeolian sands (fig. 14). thickness. the formation is c. 16 m thick at the outcrops along the coastal cliff. the formation may reach fig. 39. mud-dominated heteroliths in the lower vendsyssel formation. the light-coloured silt to very fine-grained sand beds show grading and wave ripple cross-lamination. this unit was formerly referred to as the younger yoldia clay (jessen 1918, 1931). photograph: september 2004. fig. 40. sand-rich heteroliths in the upper vendsyssel formation showing wave ripple cross-lamination. this sandrich heterolithic unit was formerly referred to as the saxicava sand (jessen 1918, 1931). tape divisions in centimetres. photograph: september 2004. 59 a thickness of up to about 25 m in the central part of vendsyssel (see fig. 125). distribution. the flat agricultural land in the vendsyssel area, lying 10–40 m above sea level, defines the top of the vendsyssel formation, and thus can be regarded to represent the fossil seabed of the younger yoldia sea. depositional environment. the formation reflects the establishment of marine conditions in the vendsyssel area after the melting back of the scandinavian ice cap in the kattegat–vendsyssel–skagerrak region. the palaeoenvironmental development may be described in terms of six events (richard 1996): the first event is represented by the erosional unconformity formed immediately after deglaciation. the second event was a rapid transgression with the establishment of a c. 60 m deep arctic marine environment. in the third event, a high sea-level stand prevailed during deposition of the clayey mud. events four to six are stages of forced regression due to the isostatic uplift in the area, but with fluctuations due to eustatic sea-level rise. age. the age of the vendsyssel formation ranges from 17 000 to 14 500 b.p. (tables 2, 3; tauber 1966; krog & tauber 1974; knudsen 1978; abrahamsen & readman 1980; aaris-sørensen & petersen 1984; nielsen et al. 1988; richard 1996; houmark-nielsen 2003). fig. 41. at the cliff section north of mårup church, the sandy mud is often highly bioturbated; in places, the shells of the bivalve hiatella arctica are found in life position in the escape trace fossils. photograph from the middle part of the section in fig. 37. photograph: september 2004. 60 structural description of sections the rubjerg knude glaciotectonic complex is differentiated into 13 structural sections, which are named after the localities recorded in the geological crosssection of lønstrup klint by jessen (1918). the sections can be grouped into three zones within the complex: a distal zone (three sections farthest to the south), a central zone (seven sections in the middle part of the complex), and a proximal zone (three sections farthest to the north). they are named, from south (near nørre lyngby) to north (at lønstrup): ulstrup, stensnæs, martørv bakker, kramrende, brede rende, sandrende, stenstue rende, grønne rende, rubjerg knude fyr, stortorn, moserende, mårup kirke and ribjerg sections. three criteria were used for defining the sections: (1) the sections had to be bordered by marked footwall ramps, (2) each section should be characterised by uniform structural architecture, and (3) the sections had to be descriptively and geographically delimited. a good example of the first criteria is the steep thrust fault separating the sandrende section from the stenstue rende section. as an example of a uniform architecture, the grønne rende section can be mentioned, and finally the mårup kirke section includes the long barren stretch from the mårup church to ribjerg at lønstrup where the lack of geographical markers as well as characteristic footwall ramps is significant. each section is described separately, a general physiographic introduction being followed by four parts: (1) tectonic architecture, (2) sedimentary units, (3) structures, and (4) interpretation of structural development. the first descriptive part provides a general description of the macro-structures. the second part presents the sedimentary deposits, and although to some degree it repeats the lithostratigraphic descriptions of the formations (see above), the detailed observations of syntectonic sedimentation are relevant to an appreciation of the structural development in the individual sections. the third part concerns the description of mesoand mini-structures (thrust faults, folds, faults, shear zones, joints, fractures, breccias, polydiapirs etc.). the description of each section is concluded with an interpretation of the formation of the structures. this interpretation should not be confused with the overall interpretation of the dynamic development of the progressive thrust-fault deformation that follows the systematic descriptions of the sections. the organisation of the section descriptions follows the general systematics of structural geology: the description of geometry, the kinematic investigation, interpretation of the dynamics and finally the analysis can be concluded by a tectonic synthesis (dennis 1972). the reader should note that the structural elements are numbered from distal to proximal. this is a consequence of the systematic analysis; in order to obtain an overview, the reader can compare the structural description given for each section with the dynamic development presented in the latter part of this bulletin. ulstrup section the southern frontal edge of the rubjerg knude glaciotectonic complex is positioned in the ulstrup section where the undeformed foreland is exposed below the hanging-wall flat of the last displaced and most distal thrust sheet (ul01, see plate 2). one of the most interesting problems addressed in the analysis of this complex is the presence of two long thin thrust sheets that were translated southwards for about 500 m from their ramps to their present positions without complete internal disruption. the ulstrup section represents the foreland margin of the thrust-fault complex, and the two flat-lying thrust sheets extend from the southern edge of the thrust front at tvonnet rende to the steep ramp thrust separating the ulstrup and stensnæs sections (plate 1). the stensnæs ramp thrust was initially regarded as the foreland thrust (pedersen 1987), but new outcrops of the southern ulstrup thrust exposed in 1996 and 1997 clearly demonstrated additional details of the frontal thrusting. consequently, the long cliff section showing horizontal bedding that had previously been regarded as a primary, undeformed sedimentary unit (fig. 42) is now interpreted as a displaced thrust sheet (ul01). the ulstrup section is truncated by a broadly horizontal glaciotectonic and erosional unconformity above which aeolian sand was deposited, either as sheet sands or as small dunes, up to 10 m high. 61 tectonic architecture the ulstrup section comprises the two flat-lying thrust sheets, ul01 and ul02. the tip line of the hangingwall ramp (the edge of the frontal thrust) is situated on the northern side of tvonnet rende (for location, see plate 1). unfortunately, the precise position is obscured by late, syntectonic erosion and sedimentation at the front of the thrust-fault complex, as well as by sand scree covering the outcrops at tvonnet rende. the northern boundary of the section is the footwall ramp and flat of the ul02 thrust sheet, which forms a transitional zone of imbricate thrusting related to the frontal part of the stensnæs section. the frontal part of the ul01 thrust sheet was displaced from the upper footwall ramp (fig. 43) along an upper footwall flat on the top surface of the foreland. at the tip of the thrust sheet, the thrust fault dips gently towards the foreland, and a small foreland-dipping ramp is also located at point 6040 m in the crosssection (plate 1, see fig. 50). the ramps probably formed due to erosion in front of the propagating thrust-fault edge. in the central part of the ul01 thrust sheet, a synform structure is associated with a chaotic breccia, interpreted as the collapse of a frost mound or sandmud diapir below the thrust fault; the synformal depression is referred to hereafter as ulstrup rende (for location, see plate 1). the trailing end of the thrust sheet starts at the upper footwall ramp of the foreland from where the hanging-wall flat is inferred to continue along the footwall flat to the footwall ramp at its trailing end. the total length of the thrust sheet is about 750 m and the displacement is estimated at 350 m, which is the distance from the footwall ramp at point 5800 m in the cross-section to the frontal termination in tvonnet rende (plate 1). the thickness of the thrust sheet varies from 10 to 20 m, decreasing towards the tip to the south, and increasing in thickness where syntectonic deposits fill the piggyback basin on the back of the thrust sheet. the thickness of the lønstrup klint formation is only 6 m in the southern part of the piggyback basin due to erosion related to elevation during thrust faulting. the ul02 thrust sheet is 600 m long, with the frontal edge situated on the northern side of the ul01 piggyback basin at point 5900 m, and the trailing end disappearing into the décollement zone at point 5300 m. the thrust fault consists of a more than 400 m long footwall flat on top of ul01 extending from the footwall ramp hinge at point 5360 m southwards to the fig. 42. the steep sandy cliff of the ulstrup section displays horizontal bedding of an apparently undisturbed deposit. however, structural analysis of thrust-fault relationships to the south indicates that it is a long flat thrust sheet displaced more than 500 m towards the foreland to the south. the cliff section is 25 m high and the view is towards the south. photograph: june 1984. 62 gently dipping frontal bend at point 5780 m in the cross-section (plate 1). the frontal bend corresponds to the hinterland-dipping limb of the flat-topped hanging-wall anticline of ul01. this hanging-wall anticline compares well with the structure of the thrust-fault model in fig. 6, and the southernmost c. 100 m of ul02 can be regarded to have been emplaced piggyback on ul01 during the translation over the frontal footwall ramp. the ul02 thrust sheet is only 5–6 m thick above the flat-topped anticline, whereas the thickness increases to 20 m at the trailing-end ramp. the displacement along the northern thrust fault is 550 m, with an uncertainty of 10–25 m depending on the interpretation of the shape of the trailing-end ramp and the amount of erosion of the frontal part at the piggyback basin at ulstrup rende. sedimentary units the sedimentary units in the ulstrup section comprise the upper part of the lønstrup klint formation, erosional remnants of the lower part of the rubjerg knude formation, and a variety of intercalations of the rubjerg knude formation distinguished here as the ulstrup beds. the l/r-unconformity between the lønstrup klint and rubjerg knude formations can be traced along the upper part of the ul01 thrust sheet in which it also forms the base of the piggyback basin. a few younger erosional unconformities, below and above the ulstrup beds, are of only local significance within the ulstrup section. lønstrup klint formation the mud-dominated lønstrup klint formation forms the main part of the thrust sheets in the ulstrup section, and has an average thickness of about 10 m (fig. 19). in the southern thrust sheet, the fine-grained sand beds are thin and only small-scale current ripples have been observed. a combination of load structures (balland-pillow) and water-escape structures (convolution and small-scale diapirs) are developed at certain horizons. above the hanging-wall flat, a zone about 2 m thick takes the form of a mobilised mud breccia, which can be characterised as a sole thrust zone. this breccia is superimposed by beds affected by a brittle type of brecciation, forming cracks and joints in an up to 4 m thick zone in the lower part of the thrust sheet. fig. 43. the upper footwall ramp of the foreland (foreland fwr) along which the ul01 thrust fault propagated (ul01hwf: ulstrup thrust sheet 01 hanging-wall flat), and from where it continued for more than 300 m over the footwall flat of the foreland. the clif f section is 25 m high and south is to the right. photograph: may 1998. 63 the northern thrust sheet (ul02) displays a more sandy part of the lønstrup klint formation (fig. 25). here the sand beds are up to 1 m thick and waterescape structures, convolute bedding and flame structures commonly disturb the primary bedding. balland-pillow structures are more common towards the trailing end of the thrust sheets. rubjerg knude formation in the ulstrup section, the rubjerg knude formation comprises three different depositional units: the main ‘background’ sedimentation of outwash sand, the glaciolacustrine ulstrup beds (fig. 25), and the coarsegrained glaciofluvial ulstrup beds (fig. 20). the main depositional unit is fineto mediumgrained meltwater sand represented in the footwall block of the foreland (fig. 43). small-scale current ripples occur in the parallel bedded sand, which is interlayered with c. 0.5 m thick trough cross-stratified beds. the glaciolacustrine ulstrup beds form a 3–5 m thick unit that is only found on the back of the ul02 thrust sheet (fig. 25). this unit consists of dark bluish grey, laminated clayey mud interlayered with a few sandy beds up to 0.5 m in thickness. the unit was deposited on a bedding-parallel unconformity, which is only discordantly developed in the northernmost trailing part of thrust sheet ul02. the unit thins out towards the south, and disappears near the hinterland-dipping limb of the flat-topped hanging-wall anticline formed above the footwall ramp of the foreland. intraformafig. 44. the mud-mobilised thrust-zone breccia consists of structureless mud with scattered clasts floating in the matrix. dilation cracks filled with sand are superimposed on the mud-mobilised brecciation fabric. this reflects two phase of cataclastic deformation: a first phase of water-over-pressured brecciation (hydrodynamic brecciation), and a second phase of brittle fracturing when the mud was consolidated, dehydrated or perhaps frozen. handle of spade is 12 cm. photograph: june 1997. tional hydrodynamic brecciation, including small-scale diapirs and slump-like features, deformed the clayey mud; such deformation is not seen in the lønstrup klint formation below the unit. the glaciofluvial ulstrup beds occur in the ulstrup rende depression between points 5900 and 6000 m in the cross-section (plate 1). this unit is a c. 8 m thick succession of meltwater gravel fining up into coarse-grained sand (fig. 20). large-scale trough crossbedding dominates the succession and clasts up to boulder size occur in the lowermost 2 m (fig. 28). lithologically, the clasts are dominated by flint, but clasts of fossil frozen sand are abundant (fig. 29). the glaciofluvial ulstrup beds are overlain by c. 5 m of mediumto coarse-grained sand of the rubjerg knude formation. on the north side of the ulstrup rende depression, between points 5900 and 6000 m, slump-folded sand beds and sedimentary breccias occur in the upper part of this succession, suggesting gravity gliding down the steep slope of the depression (piggyback basin of ul01). structures and breccias the most important structures related to the thrust faults in the ulstrup section are the breccias occurring in the thrust zones above the thrust-fault surfaces. they appear to have formed by collapse of the thrust sheet during translation. the low-angle anastomosing faults that developed in the most distal part of the thrustfault complex originated similarly during translation 64 and are associated with southerly dipping normal faults. a significant collapse structure that formed beneath the ulstrup rende depression is also worthy of note. thrust-zone breccias the thrust-zone breccias occur above the hangingwall flat of the ul01 and ul02 thrust sheets, where they affect the mud-dominated lønstrup klint formation. the thrust zone is up to 4 m thick in the most distal part of the thrust-fault system (southern part of ul01), and decreases in thickness to 1 m northwards; it can be traced along the hanging-wall flat of ul02 for a considerable distance. the thrust-zone breccia consists of mobilised mud with irregular clasts of mud scattered throughout (fig. 44). some patches may be more sandy and others more clayey, and lenses and layers of sand may be present. the mobilisation was apparently initiated as sandy mud-fluid that developed at the thrust-fault surface and extended up into the sedimentary unit (fig. 45). in many cases, the initial hydrodynamic brecciation of the thrust zone left segments along the displacement surface of the thrust fault, which were modified and developed into elongated cataclasts along the sole of the breccia zone. convolute bedding and small-scale diapirism are also present. the mobilised mud was subsequently transected by dilation cracks and sand-filled fissures (fig. 46). the dilation cracks form an irregular network and the sand-fill was injected into consolidated mud (figs 44, 46, 47). more or less horizontal sand-filled cracks have been observed in the frontal part of the ul01 thrust sheet, where they are up to 15 cm thick and appear up to 1 m above the hanging-wall flat. the sand in the cracks shows planar horizontal lamination and small-scale current ripples and a few vertical sandfilled pipes extend upwards from the cracks (fig. 48). towards the frontal tip of the ul01 thrust sheet, an increasing number of bedding parallel or low-angle anastomosing fractures and small-scale faults appear to be related to an increased rate of internal gliding. this is an indication of how close the thrust sheet was to disintegration and a loss of cohesion. zones 0.2–0.5 m thick, grading into mobilised mud, occur in between the anastomosing fractures, resulting in the destruction of bedding (fig. 49). foreland-dipping hanging-wall flat faults in the frontal part of the ul01 thrust zone, forelanddipping faults become increasingly common. these faults are either foreland-dipping (20–30°s) ramps formed by the hanging-wall flat scouring into the footwall flat (fig. 50), or sets of 50°s dipping normal faults with displacements of about 10 cm. these structures are considered to be the result of partial collapse of the tip of the foreland-dipping limb of the hangingwall ramp above a low-angle hanging-wall ramp translated along the footwall flat of the foreland. intrusive contact thrust-fault surface fig. 45. mud mobilisation along the hanging-wall flat of the ul02 thrust sheet. from the thrust-fault sur face, a sandy mud fluid intruded along fractures and up into the lower part of the thrust sheet where it formed a mud-intrusion. during the initial hydrodynamic brecciation of the thrust zone, small relict segments remained at the thrust plane where they were modified and developed into elongate cataclasts in the sole of the thrust breccia zone. photograph: july 1998; matchbox for scale. 65 fig. 46. subsequent to mud mobilisation, the thrust zone was transected by dilation cracks and fissures, which form an irregular network into which sand was ‘injected’. the mobilised mud had clearly become consolidated before the sand-filled cracks formed. the close-up inset illustrates the ‘reverse’ drag of the sand-fill (arrowed) indicating an upward direction of flow in the fissure. photograph: may 1998. 66 collapse structure in the ulstrup rende structures in the central part of ulstrup rende (figs 51, 52) are interpreted to represent a collapsed diapir. the early phase structures include thinning of the lønstrup klint formation in the thrust sheet, and formation of concave troughs in its surface. vertically or steeply dipping sand breccias with upward directed flow structures cross-cut the thrust-zone breccia. the appearance of structureless sand pockets indicates sandfill of mobilised sediment from an over-pressured zone in the subsurface. the complex of breccias and reorientated bedding is interpreted as a collapse structure; it is considered to be responsible for the formation of the ulstrup rende depression, and the disruption of the thrust sheet along steeply dipping fractures. in the breccia zone, steeply dipping sand-filled cracks and normal faults formed prior to the continued deposition of the rubjerg knude formation in the depression. on the southern side of the ulstrup rende depression, the muddy part of the thrust sheet is displaced by steeply dipping normal faults. downthrow is to the south, synthetic towards the depression, and the faults are thought to be related to the collapse of the diapir structure. interpretation of structural development there are two reasons why the ulstrup section deserves special attention. the first is that it represents a foreland thrust section with long lateral transport of thin thrust sheets, the nature of which has not previously been documented. secondly, it demonstrates the likely development of the initial stages of deformation, which the remainder of the thin-skinned thrust faulting in the rubjerg knude glaciotectonic complex also experienced before the uppermost part was eroded. thus the first phase of thrust-fault deformation is preserved here whereas it is almost never represented in the thrust-fault sections that have been fig. 47. the formation of thrust-zone breccias in the distal part of the thrust-fault complex is here illustrated in four stages of development. (1) the initial undeformed sediment (lønstrup klint formation) comprises clayey mud interlayered with thin sand beds. (2) a mud-fluid is formed above the thrust-fault surface (line with open triangles ) from where it is injected up into the layers above (see fig. 45). (3) increasing mud-mobilisation results in the formation of a structureless matrix with dispersed matrixsupported clasts of the primary sediment (see fig. 44). note the small normal faults indicating an on-going process of collapse. (4) the mobilised mud becomes consolidated and the thrust-zone breccia develops into a more brittle stage; dilation cracks form into which water-saturated sand is injected (see fig. 46). 67 more intensely deformed. the interpretation of the structural development can be summarised in the following nine stages. 1. initial thrust-fault fracturing and thrust-fault propagation took place during mobilisation of mud along the hanging-wall flat. at this stage, the thrust-zone breccia was formed due to high pore-water pressure in an unfrozen stage. 2. the ul01 and ul02 thrust sheets probably started to move along the décollement zone as one coherent thrust sheet, and first separated into two thrust sheets after the frontal part of the sheet had passed the most distal foreland footwall ramp. 3. the ramping up of the northern ul02 thrust sheet probably increased the pore-water pressure, and when this increase also affected the frontal part of ul01, diapirism was initiated under the central part of the southern thrust sheet. 4. the diapiric uplift and erosion took place in the elevated surface. this erosion extended through the rubjerg knude formation to locally intersect the l/r-unconformity. the residual coarse clastic gravel fig. 48. a subhorizontal sand-filled fracture occurring in the frontal part of the ul01 thrust sheet. the sand-filled fracture appears about 1 m above the hanging-wall flat; the sand shows planar horizontal lamination and current ripple cross-lamination. the sand-filled fracture is interpreted to have formed during ground-frozen conditions, whereas the vertical sand-filled pipe probably reflects loading of the thrust sheet when it ultimately lost its carrying pore-water pressure and settled on its hanging-wall flat. photograph: june 1997. fig. 49. subhorizontal anastomosing faulting (centre left) with mud-mobilisation developed in domains between fault fractures in the frontal part of the ul01 thrust sheet. photograph: june 1997. 68 fig. 50. a: foreland-dipping ramping of hanging-wall flat formed by scouring-erosion of the footwall flat into the top surface of the foreland. note in the close-up (b) that some hydrodynamic brecciation occurred in the footwall ramp just below the thrust zone. photograph: june 1997. on the footwall flat at the surface of the foreland in the ulstrup section. 7. the consequence of the diapir collapse was the formation of the depression in ulstrup rende. redeposited coarse-grained clastic material filled the depression, generating the glaciofluvial ulstrup beds. the occurrence of fossil frozen-sand clasts implies that part of the surface, the rubjerg knude formation, was ground-frozen – probably that part of the thrust sheet that had been elevated due to the propagation up over the central ramp. the ground-frozen condition was probably also responsible for freezing of the mobilised mud in the thrust zone and the subsequent development of sand-filled dilation cracks (fig. 47) bed on the unconformity surface was probably the original source for the large amount of coarse-grained material in the ulstrup rende depression. 5. the mud-sand volcano broke through to the surface, and the mobilised mud and sand were extruded with the release of the high pore-water pressure. 6. the surface of the diapir collapsed and gravel and sand filled the fractured structure. the collapse was probably contemporaneous with the loss of pore pressure throughout the thrust zone. the release of over-pressure in the thrust zone resulted in the final settling of the thrust sheet 69 8. the depression was ultimately filled by the sand of the upper unit of the rubjerg knude formation. the frontal edge of the northern thrust sheet propagated towards the northern side of ulstrup rende and parts of its leading tip slumped down the steep slope of the depression. this indicates the sequential and later movement of the northern thrust sheet. the depression can be interpreted in part as a piggyback basin that formed according to the model demonstrated in fig. 7. 9. finally, the uppermost sediments of the rubjerg knude formation covered the section before the glacier advanced across the area. the uppermost metre of sand was transformed into a glaciofluvialsand-glacitectonite. fig. 51. the ulstrup rende depression is interpreted to represent a combination of a piggyback basin (according to the model in fig. 7) and the collapse of an underlying mud diapir or frost-mound feature. note the steep normal fault on the left side of the depression (to the north). photograph: june 1997; spade (lower centre) for scale. fig. 52. chaotic sand/gravel breccia in the centre of the ulstrup rende depression, which is interpreted as the result of the collapse of the diapiric structure created in the subsurface below the ul01 thrust sheet. photograph: may 1998. 70 stensnæs section the stensnæs section is named after stensnæs, which is a minor point (danish: sten = stone; næs (pynt) = point) at a gentle bend in the cliff section. stones and erratic blocks, probably derived from the up to 2 m thick sandy till and the erosional unconformity below the vendsyssel formation, were formerly abundant on this part of the beach. the sandy till thins out southwards where a glaciotectonic unconformity truncates the section. the stensnæs section displays the most spectacular folds in the rubjerg knude glaciotectonic complex (fig. 53). the folds are situated at the transition from the flat-lying beds to the south (the ulstrup section) and the main thrust-fault imbrications to the north. the fold complex is truncated by an erosional unconformity forming a depression in which slump slides and sedimentary breccias derived from the tip of the thrust sheet were deposited after collapse and gravity gliding. in general, the fold complex is well exposed, whereas the transition further southwards is often covered by sand scree. during the years of study of the rubjerg knude glaciotectonic complex by the author, variations in the degree of exposure of the stensnæs section have contributed to a fuller understanding of the structural development of the section that represents the foreland margin of the thrust-fault complex. tectonic architecture the stensnæs section comprises four thrust sheets (sn01, sn02, sn03 and sn04, plate 2). the southern boundary of the section is defined by the footwall ramp of ul02, although the trailing end of ul02 is here included in the description of the imbricate duplex that hosts the fold complex. the northern boundary is the hanging-wall ramp of mb01/mb02, that is thrust up along the footwall block of sn04. the thrust sheets comprise the uppermost part of the lønstrup klint formation and a relatively thin cover of the overlying rubjerg knude formation sediments. two discrete piggyback basins (early and late) were formed above the sn01 and sn02 thrust sheets. the sn01 thrust sheet is about 30 m thick, and although the thrust sheet includes a number of small duplex imbricates it can be subdivided into upper and lower segments. the lower segment is thrust onto the ul02 footwall block with a displacement of about 30 m. the displacement of the upper segment is partitioned fig. 53. the fold complex developed in the transitional imbricate zone between ulstrup and stensnæs sections. note that the thrust faults acting as flexural slip surfaces in the folding continue into the bedding-parallel thrust-fault flats to the south (right). photograph: july 2000. 71 into a series of minor dif ferential displacements ranging from minor flexural slips along bedding surfaces in the fold structures to imbricate offsets of about 1–3 m. in the frontal part of the sn02 thrust sheet, the lønstrup klint formation is only a few metres thick, and is overlain by 10 m of sediments of the rubjerg knude formation resting on the l/r-unconformity. at the trailing end of the thrust sheet, the lønstrup klint formation is more than 15 m thick, whereas the rubjerg knude formation is cut off by the footwall ramp of sn02 (corresponding to the hanging-wall thrust fault for the imbricates of sn03). the accumulated displacement of sn02 amounts to 88 m, including 15 m up over the footwall ramp and 73 m along the footwall flat of sn01. the sn03 thrust sheet may be divided into four imbricate segments thrust onto the footwall ramp of sn02. the displacement along the hanging-wall thrust fault is 45 m, but it was also carried piggyback on sn02 during the translation along the décollement zone, which gives an accumulated displacement of 163 m for the sn03 thrust sheet. fig. 54. ball-and-pillow structures in the lønstrup klint formation in the stensnæs section. note that the size of the structures reflects the thickness of the sand beds involved. photograph: july 2000; staff divisions are 20 cm. fig. 55. dish-and-pillar structures developed in the lowermost sandy bed above the hanging-wall flat of the sn2 thrust sheet in the stensnæs section. this type of water-escape structure is interpreted to have formed during the thrust-fault translation along a hangingwall flat due to the high water pressure released from the sole of the thrust sheet. photograph: july 1998. 72 the sn04 thrust sheet is about 20 m thick, including an up to 8 m thick unit of the rubjerg knude formation resting on the l/r-unconformity above the lønstrup klint formation; this unit increases in thickness to about 15 m towards the trailing end of sn04. total displacement is about 45 m, including a hanging-wall ramp-and-flat propagation along the footwall ramp of sn03. sedimentary units the sedimentary units that crop out in the stensnæs section are dominated by the upper sandy parts of the lønstrup klint formation. a lens of a glaciolacustrine diamictite is preserved as an imbricate thrust sheet, c. 3.5 m thick and 22 m long, and is tentatively interpreted to form part of the rubjerg knude formation. the latter is mainly represented by a 6–8 m thick succession including the residual gravel deposited on the uneven surface of the l/r-unconformity. an upper unconformity truncates the sn01 and sn02 thrust sheets, and coarse gravel beds as well as olistoliths of the lønstrup klint formation were deposited in a late piggyback basin (snstru in plate 2). lønstrup klint formation the sedimentology of the lønstrup klint formation in the stensnæs section is similar to that described for the ulstrup section. the maximum thickness of the formation is about 15 m, and lithologies are dominated by 0.2–0.8 m thick fine-grained sand beds with climbing ripple lamination. thin laminae of mud and small amounts of detrital organic material commonly drape the ripples. the sand beds are interlayered with thin laminated dark grey mud beds; micro-faulting is evident in this facies. ball-and-pillow and convolute structures are abundant (fig. 54). in the uppermost 4 m of the formation, the sandy beds become thinner and the uppermost part is dominated by mud. in sand beds inferred to be situated above the hanging-wall flat, dish-and-pillar structures have been observed (fig. 55) that are interpreted to have formed by water-escape processes related to the thrusting. cheel & rust (1986) provided a model for the development of water-escape structures in glacial outwash deposits from ottawa, canada. in their model, a sequential development from simple load structures through detached ball-and-pillow structures to dish structures is demonstrated. the model predicts a strafig. 56. the piggyback basin situated above the erosional truncation of thrust sheet ul02 and sn01 in the stensnæs section (snstru in plate 2). note that normal listric faults (arrow) may be traced to the trailing end of the large olistoliths (green) deposited in the meltwater sand and gravel. this is interpreted as the result of gravity gliding of the tip of a thrust sheet, which during propagation from the north collapsed subsequent to thrust faulting up above the mean level of sedimentation. photograph: july 2000; thrusts indicated in red, unconformity indicated in purple . 73 tified distribution with convolute stratification at the base of a bed, ball-and-pillow structures dominating the main part, and dish structures formed in the uppermost part of the bed (or unit) resulting from excess pore-water fluid pressure. the observed waterescape structures in the stensnæs section as well as in the sections further to the north compare well with the model of cheel & rust (1986) (figs 54, 55; see also fig. 77), although they suggested the triggering mechanism to be earthquakes or movements due to melting of dead-ice. rubjerg knude formation in the stensnæs section, two units of the rubjerg knude formation are distinguished: (1) a lower unit dominated by fineto medium-grained glaciofluvial sand, and (2) an upper unit of varied sedimentary breccias and coarse-grained clastic deposits that infills the piggyback basin on top of the truncated thrust sheets in the section (fig. 56). the sand deposited in the lower unit may be planar parallel stratified, or exhibit large-scale trough cross-bedding with shallow troughs, only 0.5 m deep. the lower unit of the rubjerg knude formation is estimated to be 6–8 m thick, and although strongly affected by hydrodynamic brecciation, the unit compares well with the description of the lower part of the formation provided by sadolin et al. (1997). the upper unit was deposited on the erosional unconformity truncating the back of the ul02 thrust sheet, the fold and imbricate complex of sn01, and the tip of the sn02 thrust sheet (snstru in plate 2). the central part of the basin is about 7 m thick, decreasing towards both the north and south forming a relatively narrow trough. the sediments in the basin consist of large-scale irregularly trough cross-bedded glaciofluvial sand and gravel. blocks of sandy mud, which can be identified as derived from the lønstrup klint formation, were deposited as sedimentary breccias in the basin. the clasts are up to 1 × 5 m in size; blocks and clasts less than one metre in size were commonly rotated during redeposition. normal listric faults in the northern part of the basin relate to the deposition of the largest olistoliths (fig. 56). the southern part of the basin was tilted subsequent to deposition due to the fault-bend folding of the hanging-wall flat of ul02 when it propagated along the footwall flat during the latest phase of thrust faulting. thus the upper unit of the rubjerg knude formation in the stensnæs section is interpreted to have been deposited in a piggyback basin in which thrust-sheet tips thrust up from the north collapsed and gravity-glided out into the basin. structures two types of structures in the stensnæs section are related to thrust faulting: (1) the folding related to duplex imbricates, and (2) extensional faults related to push-from-the-rear in the trailing end of a thrust sheet. imbricate duplex folding the folds in the stensnæs section can be described as flexural slip folds, and have amplitudes of 1–3 m and wavelengths of 2–5 m (figs 53, 57, 58). the folds are very irregular in shape, however, and cannot be explained in terms of simple compression. analysis of the flexural slip surfaces shows that the folded layers were separated into segments and that discordant relationships exist between beds in neighbouring segments. by defining the segments as small imbricate thrust sheets in a duplex, thrust-fault terminology can be applied and hanging-wall and footwall thrusts of the individual duplex segments defined. in fig. 57 this has been done by identifying the footwall ramps (fwr), thus distinguishing five imbricate thrust sheets about 1 m in thickness. the folds in the imbricate duplex segments include both hanging-wall anticlines and footwall synclines. as documented by the refolding of the upper hanging-wall anticline in fig. 57, the folds are superimposed by sequential phases of folding and thus also phases of imbricate thrusting. figures 58 and 59 illustrate the sequential development of imbricate thrusting; the imbrication steps forward towards the footwall ramp of ul02 to the south. note also that the hanging-wall flat of each imbricate continues into an intraformational bedding-parallel thrust fault. this can be difficult to recognise in an isolated exposure, where the flats cannot be traced back to the ramp structures in the imbricate duplex (fig. 53). extensional faults in the trailing part of the sn03 thrust sheet, listric extensional faults have been observed (fig. 60). displacements along the faults are up to 0.5 m and the 74 fig. 57. flexural slip folding in the sandy beds of the lønstrup klint formation in the stensnæs section. numbers (1–4) refer to the sequential development of the thrust faults in the duplex structure, which probably formed during the collapse of the footwall ramp to the south (ul02fwr). note that the ramp thrust faults propagate into bedding-parallel flats. photograph: may 1996. fig. 58. detail of the fold structures formed by superimposed ramp propagation (see fig. 57). the thrust-fault segments in the imbricate duplex are marked by black lines and the recognised footwall ramps are annotated fwr. hanging-wall anticlines are dominant, but footwall synclines also add to the fold framework. photograph: may 1996. 75 fig. 59. simplified model of superimposed folding formed by sequential imbrication in a duplex complex. step 1 shows the undeformed beds with the ramps bordering the segments indicated. the numbers refer to the sequentially propagating hangingwall ramps. step 2 shows the first hanging-wall anticline to be formed by the progressive collapse of the trailing end of the footwall block (footwall ramp collapse). step 3 illustrates the formation of an antiformal stack during the progressive superimposed deformation. step 4 demonstrates the developed stage of superimposed fault-bend folding comparable to the structures illustrated in figs 57, 58. 76 fig. 60. low-angle listric extensional faults in the sn03 thrust sheet, in the central part of the stensnæs complex. a: an overview of the macroscopic structure, where the sn04 thrust sheet ramps up along the footwall flat of sn03 and pushes it laterally in the back. the resultant extensional fault imbricates form a boudinage-like network. b: detail of the listric extensional faults (arrows) interpreted to have been formed by push-from-the-rear. photograph: july 1996; staf f divisions are 20 cm. 77 fig. 61. illustration of the difference between the mild deformation af fecting the top of the footwall block and the strong deformation of the hanging-wall block. only a 20 cm thick zone below the shear-laminated thrust-fault surface was affected by low-angle extensional faulting grading down into a minor normal fault network. it is thus evident that the elevated water pressures supporting the thrust sheet were transmitted to the hanging-wall flat, where intense hydrodynamic brecciation took place. photograph: july 1998. 78 spacing between the faults is 0.2–0.6 m, which creates a boudinage-like network. the structures are interpreted to have been formed by push and loading of a thrust sheet ramping the formerly monoclinal fault-bend-folded thrust sheet, which responded to the gravity spreading by displacements along the extensional faults. similar mini-scale extensional normal faults are recognised in the footwall block below the hanging-wall flat of sn03 (fig. 61). it is remarkable that the thrustfault deformation only weakly affects the top of the footwall block, while the hanging-wall block is strongly affected by hydrodynamic brecciation. a zone only 20 cm thick below the shear-laminated thrust-fault surface is affected by low-angle extensional faulting, and grades down into a minor normal fault network (fig. 61). isoclinal folding has been observed in the narrow shear-laminated thrust-fault zone, adding to the impression of high strain along the thrust fault. however, it is clear that the elevated water pressures supporting the thrust sheet were located in the hanging-wall flat. interpretation of structural development the formation of the imbricate duplex fold complex in the stensnæs section is evidently related to ramp propagation. two footwall ramps are significant: the footwall ramp of ul01, which can be regarded as representing the footwall ramp of the foreland, and the footwall ramp of the trailing end of ul02. the ul01 footwall ramp acted as a stopping block for the forward push of the imbricate thrust sheets, and the propagation of this ramp was responsible for the general gentle northerly tilt of the structures. the ul02 footwall ramp was subjected to successive thrust-fault splay formation, and consequently imbrication and superimposed folding, which can be viewed as the collapse of the trailing end of the ul02 thrust sheet. the imbricate duplex fold complex of the stensnæs section can be readily compared to the model for connecting splay duplexes of the sevier thrust belt in the cordilleran fold belt (mitra & sussmann 1997). there is a close similarity with respect to the growth of the duplex by successive connecting splays of the thrust fault and the creation of folds by thrust-fault propagation. moreover, the analysis of the imbricate duplex fold complex implies that the imbrication started at the footwall ramp in the trailing end of the system and propagated towards the foreland. it can therefore be argued that the process was one of footwall ramp collapse. a simplified model for the growth of duplexes in a connecting splay duplex system is illustrated in fig. 59. with respect to the thrust-fault displacement, an active and a passive stage of translation need to be distinguished. the active translation is the amount of displacement of the thrust sheet arising from propagation along its hanging-wall thrust fault. the passive translation is the amount of transport arising from the displacement of the underlying thrust sheet which carries it piggyback fashion. during this latter translation, the underlying thrust sheet may propagate footwall ramps, which will fold the actively translating thrust sheet as well as the piggyback thrust sheets into hanging-wall anticlines. this type of deformation will create an antiformal stack. the stensnæs imbricate duplex fold complex may thus also be viewed as a mesoto macroscopic-scale antiformal stack (fig. 59). a roof thrust, which is 110 m long and about 8–12 m thick covers the sn01–sn02 duplex. the accumulated length of the three duplex segments is 175 m; c. 65 m is thus missing in the balance calculation of the southern half of the stensnæs section. it is most likely that part of the initial thrust sheet has been eroded away, but up to c. 45 m of it might have been incorporated in a foreland-dipping frontal thrust structure preserved in the chaotic imbricate fold complex. martørv bakker section the name martørv bakker is derived from the peat exposed in the coastal cliff (danish: mar = sea; tørv = peat; bakker = hills), which is covered by aeolian sand dunes above the northern end of the section. in the southern part of the section, the vendsyssel formation forms the top unit in the clif f. the vendsyssel formation was deposited on an erosional unconformity above the glaciotectonic complex. all the posttectonic deposits are prone to cliff erosion and the resultant scree partly obscures the structures in the martørv bakker section. in addition, the unconformity at the base of the planar-bedded vendsyssel formation is a focus for groundwater seep which also conceals details in the exposures. however, two important architectural elements have been recognised: (1) the common appearance of minor duplexes in the southern part of the section, and (2) the occurrence of southerly dipping normal faults in the northern part. the southern boundary of the section is the footwall ramp of sn04 in the trailing end of the stensnæs 79 section. the northern boundary is not defined by a simple reference point in the cross-section, but by the trailing end of the mb04 thrust sheet which is a combination of footwall ramp and footwall flat below the leading-edge thrust fault of the kramrende section. tectonic architecture the martørv bakker section is subdivided into four thrust sheets (mb01–mb04). the thrust sheets in the southern part of the section are subdivided into upper and lower duplex segments of which only the lower duplex segments are distinguished by separate annotations (mb02u1–mb02u3, plate 2). in the frontal part of the section, smaller imbricate duplexes are associated with syntectonically formed hydrodynamic breccias and ball-and-pillow structures. the mb01 thrust sheet is up to 20 m thick and comprises the upper part of the lønstrup klint formation. the frontal hanging-wall ramp was thrust up along the 15° dipping footwall ramp of sn04, and displacement is estimated at about 55 m. the bedding becomes steeper in the trailing end of the thrust sheet, probably due to the relatively steep dip of the ramp in the subsurface from the 20 m to the 10 m flat level. the mb02 thrust sheet is long and flat-lying and occupies more than 400 m of the section. the frontal part is only 10–15 m thick and was displaced along its upper hanging-wall flat along the footwall flat of mb01 for a distance of about 180 m. the trailing part of the mb02 thrust sheet is 20 m thick, but as indicated in the balanced cross-section it roots down to the 30 m décollement level (plate 2). the thrust sheet is displaced by a prominent normal fault in the central part of the martørv bakker section (fig. 62). above the lønstrup klint formation, a marked basin developed in the hanging-wall block of the normal fault. the fig. 62. the normal fault developed in the central part of the martørv bakker section. in the hanging-wall block to the south, the fluvial sands of the lower part of the rubjerg knude formation are preserved in a ‘fault trap’ along the fault plane. diamict sediments were deposited above the sand in a piggyback basin that developed during the thrust faulting of the mb02 thrust sheet. note that the thickness of the lønstrup klint formation in the footwall block decreases downwards along the normal fault drag. this is interpreted as a foreland-dipping limb related to a hanging-wall anticline formed prior to offset by the normal faulting. the formation of the normal fault is interpreted to be related to a foreland-dipping limb of a hanging-wall anticline at the tip of a subsurface thrust-sheet segment (mb02u3 in plate 2). photograph: june 1993. 80 sediments in this basin were described as moraine sand by jessen (1918, 1931). the mb03 thrust sheet is 30 m thick and made up of the lønstrup klint formation, which is here strongly deformed by hydrodynamic brecciation and dislocated by a number of bedding-parallel minor thrust faults. the frontal part is flat-lying, whereas the dip of the bedding increases to 25–30° at the trailing end indicating a bend over an upper hinge on top of the footwall ramp of mb02. in the exposed part of the section, the mb03 thrust sheet is only about 120 m long, and it is inferred that the foreland-dipping structures in the frontal part reflect re-orientation due to hanging-wall ramp propagation along the footwall flat of mb02. the mb04 thrust sheet forms a flat-topped hanging-wall anticline above the footwall ramp of mb03. the thrust sheet is displaced along a normal fault parallel to the foreland-dipping bend in the top of the frontal part of mb03. the normal fault does not displace the footwall flat of mb02, and is therefore regarded as a structure related only to the framework of mb03 and mb04. the structural framework in this part of the section may be characterised as an antiformal stack, including thrust sheets mb03 below and kr01 above mb04. sedimentary units the martørv bakker section is dominated by the lønstrup klint formation. however, the most important sedimentological feature in the section is related to fig. 63. hydrodynamic brecciation in the lønstrup klint formation in the northern part of the martørv bakker section. photograph: june 1993. fig. 64. slump-fold structures formed in the thin-bedded sand layers enveloped in dark muds of the diamict sediments in the piggyback basin in the martørv bakker section. photograph: october 1998. 81 the basin developed at the top of the hanging-wall block connected to the normal fault displacing the mb02 thrust sheet (fig. 62). the deposits in this basin are regarded as an exotic part of the rubjerg knude formation, and are described under this heading below. the southern part of the martørv bakker section is unconformably overlain by the vendsyssel formation, the initially glaciotectonic truncation being superimposed by a post-glacial erosional unconformity. a holocene erosional unconformity truncates the vendsyssel formation as well as the glaciotectonic unconformity, and the peat deposited on this unconformity is up to 2 m thick in the northern part of the section, where it is covered by modern aeolian dunes up to 20 m high. lønstrup klint formation the lower part of the lønstrup klint formation is mudrich. it is exposed in the northern part of the section, where it was thrust above the hanging-wall flat from the décollement surface 25–28 m below sea level. the upper part of the formation is dominated by 0.5–1.5 m thick sand beds interlayered with thin beds of horizontally laminated mud, which typically has been mobilised to form hydrodynamic breccias. in the southern part of the section, the beds are strongly affected by ball-and-pillow deformation (fig. 63). rubjerg knude formation the rubjerg knude formation comprises two units: (1) a lower 3 m thick sand unit only exposed along the prominent normal fault in the central part of the martørv bakker section, and (2) a c. 15 m thick diamictite interpreted as a glaciolacustrine mud with redeposited clasts. the first unit was deposited on the l/r-unconformity, and comprises light yellowish medium-grained sand (fig. 26). the diamictite unit rests partly on the lower sand unit, and partly on the l/runconformity at the top of the lønstrup klint formation. the diamictite is dark grey, and comprises weakly laminated mud interbedded with structureless, irregularly distributed matrix-supported beds containing fig. 65. a schematic diagram explaining the development of the hydrodynamic brecciation displayed in fig. 63. the formation of the structure was the result of three phases of deformation. in the first phase, a succession of sandy turbidites interbedded with mud (1) was affected by loading to form the ball-and-pillow structures (2). in the second phase, the ball-and-pillow structures were displaced by thrust faulting (3). during the third phase, the mobilised mud intruded up thorough the thrust-fault surface (4), demonstrating the syntectonic development of the hydrodynamic brecciation. 82 unsorted clasts in a matrix of sandy mud (fig. 26). in the lowermost 2 m of the diamictite succession, the matrix-supported clasts were probably derived by redeposition of coarse-grained material eroded from the l/r-unconformity to the north. about 3 m up in the succession, a c. 1 m thick bed occurs with clay clasts 10 cm in size deposited in a weakly clay-laminated and sand-streaked silty mud. this bed is overlain by three 1.5–3 m thick units of isoclinally slump-folded, thin-bedded, fine-grained sand encased in dark structureless mud (fig. 26). these units are interpreted as slump-folded sheets derived from the lønstrup klint formation (fig. 64). the middle and upper slump-units are separated by a c. 5 m thick interval dominated by sandy mud with scattered clasts and a few thin sand beds. the uppermost sand bed was not af fected by slumping and shows large-scale trough cross-bedding. the lower part of the diamictite succession is strongly disturbed by hydrodynamic brecciation with mesoscopic-scale diapirs rising from the top of the lønstrup klint formation and penetrating upwards into the diamict sediments. this indicates that the diamictites were part of the main sedimentation affected by glaciotectonic disturbances. structures in the martørv bakker section, three types of structural elements were studied: (1) imbricate duplexes dominating the southern part of the section, (2) the normal fault in the central part of the section, and 3) hydrodynamic brecciation contemporaneous with, or superimposed on, ball-and-pillow structures (figs 63, 65). imbricate duplexes the imbricate duplexes in the southern part of the section constitute rhomb-shaped segments 10 to 25 m in size bounded by low-angle thrust faults. minor imbricates may occur along the thrust faults, but the thrust faults are mainly narrow fracture surfaces without significant brecciation. a mini-scale example of duplex formation is shown in fig. 66. although the structure is related to intraformational deformation of the beds, it illustrates instructively the formation of hanging-wall ramp propagation of a stacked footwall ramp. thus the footwall ramp is formed by the trailing edges of two duplex segments that were displaced one over the other to form a single planar ramp for the propagation of the upper thrust sheet. flame-like upright minor anticlines are interpreted as compressed hanging-wall anticlines formed during the sequential propagation of the various ramps. on top of the upper footwall hinge, a radial flame structure probably indicates the site of incipient diapirism (fig. 66). the formation of foreland-dipping thrust structures above the tip of a lower duplex segment is also apparent. normal fault the normal fault in the central part of the section is an uneven fault plane striking e–w with a dip of about 45° to the south. displacement along the fault plane is about 25 m, and a set of minor normal listric faults displace the top of the hanging-wall block (fig. 62). the footwall block comprises the lønstrup klint formation, which decreases in thickness southwards and forms an irregularly folded drag along the fault plane. at the top of the hanging-wall block, the diamict sediments described above are discordantly superposed on the light-coloured sand at the base of the rubjerg knude formation. hydrodynamic brecciation ball-and-pillow structures occur in the sand-rich parts of the lønstrup klint formation, where hydrodynamic mud mobilisation created chaotic breccias (fig. 63). the initial size of the sand ball-and-pillow structures is related to the primary thickness of the beds, but subsequent to the sedimentary load deformation they were distorted and deformed during thrust-fault related mud remobilisation. in fig. 63, the distorted ball-andpillow structures can be seen to be displaced by minor thrust faults, and these thrust faults were intruded by mobilised mud. the sequential development of this hydrodynamic brecciation is illustrated in fig. 65, where three phases of deformation are recognised, although these probably developed progressively during thrust-fault displacement and related loading of superposed thrust sheets. interpretation of structural development the thickness of mb01 implies that the décollement surface in the southern part of the section is situated 83 at the 20 m level. the thickness of mb03 is 30 m, implying that the décollement level stepped down 10 m somewhere in the central part of the section. a lower footwall ramp and a corresponding hangingwall ramp must therefore be included in the balanced cross-section. the l/r-unconformity reference surface on top of the mb02 thrust sheet was about 20 m above present sea level prior to the normal fault displacement, which indicates that a duplex 20 m in thickness is situated below the trailing end of mb02. the lower footwall ramp responsible for the fault-propagation folding of the antiformal stack in the northern part of the section, estimated from the bend of the trailing ends of mb02 and mb03, must be situated below mb02. the structural model therefore suggests that a subsurface duplex was formed by segments of the mb02 thrust sheet situated between the 20 and 30 m levels (mb02u1–mb02u3). the footwall ramp thus constitutes two 10 m thick duplex segments stacked on top of each other. consequently, a hanging-wall anticline with a foreland-dipping limb formed above the hanging-wall ramp of mb02 and was translated along a footwall flat. the model further suggests that the hanging-wall fig. 66. a model of duplex formation is here illustrated by a mini-scale structure related to intraformational deformation of beds in the lønstrup klint formation, central part of martørv bakker section. the duplex comprises two segments, which were derived from the bed underlying the lower footwall flat (lower fwf). the footwall ramp for the segments is situated to the left outside the frame of the figure. the lower segment is a relatively short one, which was thrust over by the upper segment during the push from the ramping of the upper thrust sheet. during propagation up the footwall ramp, the trailing edges of the two segments were displaced to form one planar ramp for the upper thrust sheet, which was further translated over the duplex to a foreland-dipping bend created above the tip of the lower duplex segment. the flame-like upright anticlines are interpreted as compressed hanging-wall anticlines formed during the sequential propagation of the various ramps. note the radial flame structures at the upper footwall hinge indicating incipient diapirism. the small normal faults to the left of the trowel (15 cm in size) are thought to reflect similar forelanddipping features in the subsurface. fwh, footwall hinge; fwr, footwall ramp; ufwf, upper footwall flat; hwf, hanging-wall flat; r, ramp. photograph: june 1993. 84 structure is a composite feature partly constructed by the hanging-wall anticline related to the tip of the mb02u3 segment folded over the footwall ramp of mb02u2, and partly by the hanging-wall anticline related to the translation of the main hanging-wall ramp of mb02 along the 10 m level. the foreland-dipping limb of this structure corresponds well with a 45° south-dipping normal fault with a vertical displacement of about 20 m. it is therefore concluded that the northern slope of the diamict sedimentary basin was formed by the normal fault reflecting the foreland-dipping limb of a hanging-wall anticline. the southern more gently dipping slope of the basin was formed by the bend of mb02 due to its propagation up along the footwall ramp and flat of mb01. this footwall thrust fault is a composite imbricate duplex, which hampers the exact distinction of ramp-flat relationships. the slumpfolded units in the basin are interpreted as the result of gravity slides derived from the crest of the hanging-wall anticline or the tip of the mb04 thrust sheet propagating from the north. the sediments filling the basin represent redeposited material derived from the thrust-fault elevated part of the lønstrup klint formation, the coarse-grained clastics on the l/r-unconformity, and the lowermost part of the rubjerg knude formation. the basin is interpreted as a piggyback basin with syntectonic deposition during the translation of the mb02 thrust sheet. kramrende section from the south, the first significant macroscopic-scale diapir occurs in the kramrende section (the kramrende diapir). although mobilisation features also occur in sections farther to the south, this diapir is regarded as the most distal in the glaciotectonic thrustfault complex. the thickness of the thrust sheet hosting the kramrende diapir suggests it is related to the fig. 67. the northern part of the kr01 thrust sheet where the lithostratigraphic reference section of the lønstrup klint formation (fig. 21) is situated. the thrust sheet is fault-bend-folded up along an initially low-angle (c. 8°) footwall ramp, which was subsequently folded into the present more steeply dipping orientation. note the reverse faults interpreted as small back-thrust faults. photograph: june 1993. 85 deep level of thrust-fault rooting, which is about 30 m below the l/r-unconformity. the thrust sheet to the south of the kramrende diapir and two thrust sheets to the north are included in the section because they are all affected by the structures related to the diapir. the frontal edge of the kramrende section is formed by the footwall ramp beneath the first thrust sheet (kr01, see plates 1, 2). this thrust fault is identical with the trailing-edge footwall ramp of the martørv bakker section, which is responsible for the marked monoclinal fault-bend folding of the kr01 thrust sheet (fig. 67). the steps to the beach are situated in the gully between the kr01 thrust sheet and the kramrende diapir. the steps lead up to the summerhouse area at oddervej, and are referred to as the kramrende steps or the oddervej trappe. tectonic architecture the kramrende section consists of four thrust sheets (kr01–kr04; plates 1, 2). the frontal thrust sheet (kr01) is ramped over the mb04 footwall ramp in the martørv bakker trailing thrust sheet. at the north end of kr01, the l/r-unconformity is situated about c. 10 m a.s.l., which indicates that the kr01 hanging-wall ramp propagated along an intermediate footwall flat (fwf). the upper ramping along the mb04 footwall ramp is fig. 68. ball-and-pillow structure developed in the sandy turbidite bed between 4 and 5 m in fig. 21. the structure is interpreted as a load structure formed immediately after sedimentation. additional load structures can be seen at the base of the sand bed, where flame structures related to the underlying clayey bed intrude the base of the sand bed. above the balland-pillow structure, pinch and swell structures within the sand bed are also interpreted as gravity load structures. photograph: june 1993. fig. 69. in the upper part of the kr04 thrust sheet in the kramrende section, the rubjerg knude formation forms a piggyback basin, which is overthrust by the br01 thrust sheet in the southern part of the brede rende section. the thrust fault displayed in the photograph is a hanging-wall flat for the thrust sheet br01 (br01hwf) and footwall ramp of thrust sheet kr04 (kr04fwr). a minor satellite thrust fault was formed below the main thrust at a late stage of fault propagation after the sand of the rubjerg knude formation had been somewhat compacted. photograph: june 1984. 86 responsible for the fault-bend-fold appearance of the kr01 thrust sheet. above the frontal part of the mb04 footwall ramp and flat, the kr01 thrust sheet is folded into a flat-topped anticline. in the involute part of this anticline, a splint or horse is present (the kr01 splint). this is a small thrust-sheet wedge torn off during thrust propagation, which created peculiar anticlinal features in the structural profile. the kr02 thrust sheet takes the form of a major diapir. initially the diapir was a thrust sheet that was displaced up along the kr01 footwall ramp and above the back of the kr01 thrust sheet. the kr03 and kr04 thrust sheets situated on the back of kr02 are characterised by marked dif ferences in the thickness of the rubjerg knude formation. in the kr03 thrust sheet, the thickness is only about 8 m, whereas in kr04 the thickness of the rubjerg knude formation is up to 20 m. this indicates that the kr04 thrust sheet was thrust over the upper footwall flat of kr03 at an earlier stage compared to a probably longer time of deposition in the piggyback basin of kr04. the c. 20° northerly dipping inclination of the kr03 footwall flat and related parallel structures is due to the bend caused by the propagation of kr02 along the footwall ramp. the main décollement level below the kramrende section is situated at the 30 m level. sedimentary units the description of the sedimentary units in the kramrende section is mainly based on sedimentological fig. 70. the kr01 footwall syncline developed below the kr01 footwall ramp (kr01fwr), which is overlain by the kr02 hanging-wall flat (kr02hwf). note how the mobilised mud migrated from the steeply dipping limb of the footwall syncline up into the kramrende diapir, intrusively penetrating the thrust fault. photograph: june 1995. fig. 71. mobilised mud from the lower part of the lønstrup klint formation in the kr02 thrust sheet intruded the turbidite sand beds in the upper part of the formation. photograph: june 1995. 87 logging of the lønstrup klint formation in the kr01 thrust sheet (fig. 21). the detailed section of the rubjerg knude formation at the top of this log is uncertain due to poor exposure and dif ficulty of access. the rubjerg knude formation exhibits variations in thickness throughout the kramrende section, and the description herein is based on scattered observations. a c. 1 m thick homogeneous, structureless sandy till caps the kramrende section. no preferred clast fabric has been recognised in the till, and its stratigraphic position is uncertain, although the occurrence of rare rhomb porphyry erratics may indicate an affinity with the norwegian ice (kattegat till formation). lønstrup klint formation in the kramrende section, the lower part of the lønstrup klint formation mainly occurs in the kramrende diapir, within the kr02–03 thrust sheets. remobilisation of the mud has obliterated primary sedimentary structures, and the diapirism also affected the thrust sheet kr03, so that only primary bedding is recognisable in the uppermost part of the formation. the primary sedimentary structures of the upper part of the lønstrup klint formation are reasonably preserved in the kr01 thrust sheet (fig. 20), although the sediments here, dominantly constituting sandy turbidite beds (10– 50 cm thick), are strongly affected by hydrodynamic brecciation creating ball-and-pillow structures (fig. 68). the top of the lønstrup klint formation in the kramrende section is truncated by the l/r-unconformity, which displays an erosional relief of 1–2 m. rubjerg knude formation the rubjerg knude formation varies in thickness from only 6–8 m at the top of the kr03 thrust sheet to about 20 m in the upper part of the kr04 thrust sheet (it is absent in kr02). the main part of the formation comprises thick-bedded, large-scale cross-bedded, medium-grained light yellow-grey sand. some beds are rich in heavy mineral sand, which occurs in parallel-laminated strata. in the kr04 thrust sheet, the heavy mineral beds are present about 6 m above the l/runconformity and again about15 m above the base. these beds have a characteristic content of small (0.1– 1 cm), grey clayey mud-clasts, which are interpreted to reflect erosion of muddy thrust sheet units in the vicinity of the depocentre. pedersen (1987) described deposits, referred here to the rubjerg knude formation, that accumulated syntectonically in footwall growth synclines. the deposits were characterised as ‘banana’ shaped basins, and a similar type of sedimentary/structural feature occurs in the kr04 thrust sheet (fig. 69). the rubjerg knude formation at the top of the kramrende section can be regarded as a piggyback basin and the footwall syncline as a growth-fault syncline. at the top of the piggyback basin, large-scale trough cross-bedded sand is truncated by small satellite thrust faults (similar to that shown in fig. 69), which are truncated by superposed c. 1 m thick trough cross-bedded sand beds. three succeeding developments of this interference between thrusting and deposition reflect the syntectonic depositional dynamics of the piggyback basin. structures structures of significance in the kramrende section are described under the following headings: (1) thrust faults, and in particular associated footwall ramps and footwall synclines, (2) the kramrende diapir, with the diapiric breccias and intrusive structures formed by mobilised mud, and (3) reverse faults, here interpreted as back-thrust faults. thrust faults the kr01 thrust sheet is bounded by the hangingwall ramp and flat (kr01hwr and hwf) at the base, and the kr01 footwall ramp and flat at the top (kr01fwr and fwf). the wedge-shaped geometry of the kr01 tip implies that the kr01hwr had a low angle of inclination, dipping about 15°n. towards the trailing end of the thrust sheet, the ramp passes into a hanging-wall flat which is parallel to bedding in the lønstrup klint formation. this thrust fault now dips at 25°n, although it is a hanging-wall flat situated on a footwall flat. this is due to the fault-bend folding related to the thrusting in the trailing part of the martørv bakker section. at its trailing end, the kr01 thrust sheet is folded into a footwall syncline (fig. 70). the bend of the northern limb in the syncline lifted the l/r-unconformity up to a position nearly 5 m higher than in the horizontal involute part of the fold, and the bedding in the lønstrup klint formation was tilted into a nearly vertical position (fig. 70). 88 the kr02 thrust sheet was thrust up along a c. 25° dipping footwall ramp (kr01fwr) onto the upper footwall flat above the rubjerg knude formation of the kr01 thrust sheet. the kr02 thrust fault is apparently a hanging-wall flat which indicates a rather long displacement for thrusting. the top surface of kr02 is a footwall flat upon which the kr03 hanging-wall flat is situated, only bringing different stratigraphic levels of the lønstrup klint formation into contact. the dip of the thrust fault is parallel to the dip of the kr02– kr01 thrust fault. the footwall flat of the kr03 thrust sheet is overlain by a c. 80 m long hanging-wall flat of kr04. the lønstrup klint formation is only about 8– 10 m thick above the hanging-wall flat (kr04hwf), indicating a fairly long intermediate flat (at the 10 m level below the reference surface). at the trailing end of kr04, the thickness increases indicating the presence of a hanging-wall ramp at the base of the thrust sheet, and a footwall syncline similar to that described in kr01 is present. in the upper part of the kr04 thrust sheet, the sediments deposited in the footwall growth syncline became overturned along the northern limb during translation of the hanging-wall ramp of br01, as described above (fig. 69). kramrende diapir the kramrende diapir constitutes the main part of kr02. the diapirism also affected the trailing end of kr01 (fig. 70) as well as some parts of kr03. as shown in fig. 71, mud of the lower part of the lønstrup klint formation in kr02 became mobilised and intruded the overlying turbidite sand beds and also penetrated upwards into the overlying kr03 thrust sheet. figure 70 illustrates a thrust fault penetrated by intrusive mud at the footwall ramp of kr01. here, the mobilised mud from the steeply inclined northern limb of the footwall syncline intruded into the mud-breccia along and above the hanging-wall flat of kr02. a large part of the boundary between kr02 and kr03 was deformed in a similar way and the primary layering destroyed. reverse faults in the kramrende section, significant steeply dipping reverse faults occur in the northern part of kr01 (fig. 67), and in the middle part of the kr03 thrust sheet. the displacement is only about 30–50 cm on the steep south dipping faults in kr01; the spacing between the faults varies from 3–9 m, and often the faults can be traced down into the tectonic breccia above the fig. 72. reverse fault-splay fan developed on the back of kr03 and also displacing the overlying kr04 thrust sheet. the structure is interpreted as a back-thrust fault-splay formed during the kr03 propagation over the upper footwall ramp situated on the back of kr02. photograph: july 1994; figure at fault-splay centre for scale. 89 hanging-wall flat. in the kr03 thrust sheet, the reverse faults form a fault splay fan with individual faults dipping moderately to steeply to the south (fig. 72); displacement varies from about 20 cm up to about 3 m. the reverse faults appear to have formed during thrust-sheet propagation over a ramp hinge and are interpreted as back thrusts. a bend over a shallow dipping ramp will only result in minor displacement on steeply dipping back thrusts, whereas bending over steeply dipping ramps creates low-angle back thrusts with potentially greater displacements. interpretation of structural development the balanced cross-section indicates that the kr01 thrust sheet was about 350 m long, of which a major part of the front tip has been eroded away. the amount of displacement is deduced from a series of balanced approximations to be 160 m (see plate 2). the footwall ramp of kr01 is interpreted to root in the 30 m décollement level, which indicates that the lower hanging-wall ramp of kr01 was displaced onto the intermediate footwall flat above the trailing segment of mb04. thus the lift and steep tilt of the northern limb in the kr01 footwall syncline is interpreted to have formed during the displacement of the lower hanging-wall ramp (kr01hwr) along the trailing-end footwall flat of mb04 (mb04fwf). the tip of the kr02 thrust sheet has been eroded away. to avoid exaggeration, the thrust fault is interpreted to have continued only about 15 m up in the air further to the south, which implies that displacement of kr02 was in the order of 50 m. the main part of the kr02 thrust sheet present in the cross-section is above the hanging-wall flat brought up from the 30 m décollement level. it is evident that the mobilised mud was derived from this low level and that it was activated in diapirism during the thrust-sheet propagation over the footwall ramps (kr01fwr) and its hinge bend. it is difficult to estimate the displacement for the fig. 73. the brede rende diapir in the frontal part of the brede rende section. the frontal part of the br02 thrust sheet was thrust up on its hanging-wall ramp (br02hwr) along the footwall ramp, which turned into the kr04 footwall flat (kr04fwf). at an early stage of thrusting, probably while the hanging-wall ramp passed by a lower footwall hinge, diapirism developed. the mobilised mud also intruded the l/r-unconformity and formed mushroom-shaped diapirs in the rubjerg knude formation. photograph: june 1984. 90 kr03 thrusting. its relationship to the footwall flat of kr02 indicates that it was thrust along a hanging-wall flat in the order of 60 m. however, before the kr03 thrusting was complete, the kr04 thrust sheet was already emplaced on its back. the displacement of kr04 can be determined in the cross-section to be 66 m. the implications of kr04 being thrust onto kr03 are that the sedimentation of the rubjerg knude formation on top of kr03 ceased, and with the continued propagation of kr03 over the footwall ramp, the back-thrust splay also affected the kr04 thrust sheet that was being passively transported piggyback on kr03. brede rende section for more than a century, groundwater drainage has been concentrated at a spring at brede rende. from the spring, situated at the base of the cliff, a stream has over the years eroded a large funnel-shaped gully behind the cliff facing the sea. groundwater erosion successively stepping backwards is thus responsible for the wide gully and for the locality name (danish: brede = wide; rende = gully). the groundwater transmissivity is, of course, governed by the geology of the brede rende section, such that the spring wells out from the unconformity surface between the clayey lønstrup klint formation and the permeable sand of the rubjerg knude formation. the water initially drained in a southerly direction, but the present northerly drainage system is exposing the structures of the northern flank of brede rende in an isolated cliff. it is likely that this cliff will be completely removed by erosion by the sea as well as by the stream within the next few years leading to the formation of a broad gully at this location. important structural features currently exposed in the brede rende section comprise a polydiapiric complex, the brede rende diapir in the frontal part, the prominent brede rende normal fault (brnf) in the central part, and a series of duplexes stacked in the trailing end of the section. tectonic architecture the brede rende section comprises eight thrust sheets, annotated br01–br08 (plate 2). the southern and frontal boundary of the section is the footwall ramp of kr04. the northern boundary is the thrust fault which partly acts as the br08 footwall ramp and flat, and partly is the hanging-wall ramp for the sr01 thrust sheet in the sandrende section. the br01 thrust sheet is a relatively thin sheet that was displaced up along the kr04 footwall ramp, which is c. 30 m thick and dips about 35°n. above br01, the br02 thrust sheet was displaced more than 50 m along its hanging-wall ramp onto the upper footwall flat of kr04. the lower part of the thrust separating br01 and br02 has been destroyed by penetrating diapirism, and together with br03 these frontal thrust sheets in the brede rende section constitute the brede rende diapir (figs 73, 74). the thrust sheets can still be regarded as individual coherent elements, although their fig. 74. detail of the internal structure of the brede rende diapir. although the structure appears as a chaotic mixture of disrupted sand beds ‘floating’ in a disorganised fashion in the mobilised mud-matrix, some of the features could be interpreted as relicts of hanging-wall anticlines (see dashed lines) formed in a developed stage during thrusting up along steep ramps. photograph: june 1985. 91 boundaries and internal structure have been strongly distorted by diapirism. br03 is the longest thrust sheet in the brede rende section, when the trailing lower segment is included (see plates 1, 2). this is a constructional convention based on the consideration of which of the hangingwall ramps should be traced down to the décollement surface, and thus determine the annotation of the subsurface duplex sheets (plate 2, see later). the br04 thrust sheet is about 300 m long and is displaced by the brede rende normal fault (brnf) (fig. 75). north of the brnf, the br04 thrust sheet was thrust along an intermediate br03 footwall flat, and south of brnf the upper hanging-wall ramp and flat of br04 were thrust over the upper footwall flat of br03. the amount of displacement along thrust faults in this part of the brede rende section is 50 m (measured in the crosssection of plate 1) for br03 as well as br04. in br04, the rubjerg knude formation reaches its maximum thickness of about 20 m in the brede rende section, whereas the cover of rubjerg knude formation on the back of br02 and br03 is less than 5 m thick. the br05 thrust sheet is relatively short and located between br04 and br06. the displacement along its hanging-wall ramp is about 80–90 m and the initial ramp-angle was c. 12°. the thickness of the rubjerg knude formation on top of br05 is only about 5 m, which indicates that the thrusting of the br06 thrust sheet propagated early in the thrust development of the brede rende section. the piggyback thrusting of br06 on br05 on br04 is one of the best examples of a duplex structure in the rubjerg knude glaciotectonic complex. the br06 thrust sheet is a relative long and thin thrust sheet. to the north, the trailing end of br06 was thrust up over the footwall ramp of br05. from the footwall ramp hinge, an upper hanging-wall flat (br06hwf) was displaced along the upper br05 footfig. 75. the brede rende normal fault (brnf). the frontal part of the br06 thrust sheet has a normal displacement of about 20 m down through the 45° dip normal fault, which can be measured from the hanging-wall flat of br06 (br06hwf ) north of the normal fault to the br06hwf south of the normal fault. in the footwall block of the brnf, a series of minor normal faults make a stepwise displacement of the downthrown hanging-wall block. in the hanging-wall block of the brnf, the bend of the thrust-fault structures may be characterised as a roll-over anticline. note that the br06 hanging-wall flat transforms into a hanging-wall ramp (br06hwr), which was thrust-displaced along the upper footwall flat of the br04 thrust sheet (br04fwf). photograph: june 1984. 92 wall flat for about 200 m. as noted above, the br06hwf developed above the br05 thrust sheet at an early stage. the br06 thrust sheet is divided into two segments by the brnf (fig. 75). south of the brnf, the upper hanging-wall ramp of br06 was emplaced on the br04 upper footwall flat (fig. 75). the br07 thrust sheet was thrust piggyback onto the trailing-end segment of br05 and propagated up along the footwall ramp of br06. it has very chaotic internal structures dominated by polydiapirism. the position of the reference surface (l/r-unconformity) at an elevation of 20–30 m above sea level indicates that the br07 sheet was ramped up onto the flat above a duplex composed of the trailing segments of br03 and br05 (see later). there is only a thin cover of less than 5 m of the rubjerg knude formation on the back of br07, which is overlain by the br08 hanging-wall ramp. the br08 thrust sheet is the northernmost and uppermost sheet in the brede rende section. it is the smallest thrust sheet in the section, the thin frontal tip of the thrust sheet consists of the uppermost part fig. 76. ball-and-pillow structures superimposed by chaotic hydrodynamic brecciation in the upper part of the lønstrup klint formation in the brede rende section. the brecciation was formed by polysequential diapirism during thrusting of the br06 hangingwall flat (br06hwf ) over the hinge to the upper footwall flat of br05 (br05fwf) situated in the left side of the photograph. photograph: june 1993. 93 of the lønstrup klint formation, but with an up to 12 m thick succession of the rubjerg knude formation on top of the l/r-unconformity. due to the bend up along the br06 footwall ramp, the inclination of the br07 and br08 thrust sheets is c. 25°n. sedimentary units the lønstrup klint formation is strongly affected by ball-and-pillow load structures and hydrodynamic brecciation. the maximum thickness of the formation exposed is only about 20 m (tentatively measured in br02). the thickness of the rubjerg knude formation varies from thrust sheet to thrust sheet, indicating differential thrust-fault movement that either closed the piggyback sedimentation and/or lifted the formation up to a position exposed to erosion. the rubjerg knude formation was also subjected to hydrodynamic brecciation. at the top of the central part of the brede rende section, a glacitectonite and associated glaciotectonic imbrications are interpreted to be related to the advance of the norwegian ice; the sandy till is interpreted to be the kattegat till formation. lønstrup klint formation the lower and intermediate parts of the lønstrup klint formation are characterised by dark clayey mud. the interval 5 to 10 m below the l/r-unconformity is dominated by a few thick beds of light coloured sandy turbidites, and the uppermost 5 m is formed by thinbedded sand beds interbedded with mud. the size of fig. 77. small-scale ball-and-pillow structures distorted and intruded by water-escape injection. note that dish structures were formed above the water-escape pipe. detail of chaotic brecciation in the upper part of the lønstrup klint formation; frontal part of the br06 thrust sheet in the brede rende section. photograph: june 1998. fig. 78. small-scale disharmonic undulations formed by polysequential diapirism in the thinly interbedded clays, silts and finegrained sands of the upper part of the lønstrup klint formation. note the fold accentuation of the climbing ripple lamination in the central part of the figure. frontal part of the br06 thrust sheet in the brede rende section. photograph: june 1998. 94 ball-and-pillow structures is typically related to the initial thickness of the sand beds, and the subsequent hydrodynamic brecciation and chaotic structures formed during water-escape activities (figs 68, 76–78). the l/r-unconformity at the top of br04 truncates a large ball-and-pillow structure at the top of the lønstrup formation just north of the normal fault (fig. 79). this implies that some of the load structures, and possibly also initial water-escape dynamics, had commenced prior to the development of the unconformity. it may be that this phase of ball-and-pillow formation was initiated by the drainage of the large lake basin (see sadolin et al. 1997). thus the initiation of balland-pillow formation can be viewed as the consequence of vibration created by an increased water transport over the beds. at this locality, the formation of ball-and-pillow structures was clearly not the effect of loading by over-thrusting, but only the result of density variation of the primary sedimentary layers, since the top of the lønstrup klint formation was undergoing erosion and the gravel bed on the l/runconformity was deposited subsequently. rubjerg knude formation the rubjerg knude formation reaches a thickness of 15 m in the upper part of the br04 thrust sheet, but in the rest of the brede rende section it is less than 10 m thick. the relatively thin nature of the formation (3–5 m) in br02, br03, br05 and br07 is interpreted to indicate that these basins were over-thrust or thrustelevated at an early stage of thrust propagation. in contrast, deposition persisted in the piggyback basins of br04 and br06 before their upper footwall flats were overthrusted and deposition ceased in the basins. structures three types of structural features are described from the brede rende section: (1) diapir structures, including mesoscopic-scale sequential polydiapirs and hydrodynamic brecciation, (2) the brede rende normal fault (brnf), and (3) frost wedges. fig. 79. a large ball-and-pillow structure in the upper part of the lønstrup klint formation, truncated by the l/r-unconformity. this relationship demonstrates that at least part of the loading occurred prior to the thrust-fault emplacement. photograph: june 1998. facing page – upper: fig. 80. normal fault network in the footwall block of the brede rende normal fault developed in the br04 thrust sheet. photograph: june 1998. facing page – lower: fig. 81. frost wedges recognised in the rubjerg knude formation. the one on the right side of the spade (a) has well-developed, ‘upwards-fanning’ small-scale normal faults, whereas the one to the left of the spade (b) is a 5–10 cm wide fracture with a sand-fill. photograph: june 1997. 95 96 diapir structures the term diapir as used here follows the definition of weinberg & schmeling (1992 p. 425): “diapir is the non-genetic geological term applied to ductile intrusive structures. many diapirs may develop due to rise of gravitationally unstable buoyant fluids through denser overburden. such gravitationally unstable configurations consisting of viscous layers are known as rayleigh-taylor instabilities.” in the brede rende section, the diapir structures can be divided into two types: (1) smallto medium-scale diapirs that developed into hydrodynamic breccias in which primary sedimentary lamination is locally preserved, although distorted and irregularly folded, and (2) mediumto large-scale diapirs where mobilised mud intrudes overlying stratigraphic levels or thrust units. the first type of diapirism corresponds to the sequential polydiapirs of weinberg & schmeling (1992). these are initiated as small undulations or even flame structures, that develop into irregular upright folds with numerous minor undulations on their flanks (figs 76, 78). when the viscous mud broke through the bedding it formed intrusive pipes (fig. 77), and either spread out laterally between layers or released water, forming dish-and-pillar structures in the overlying beds fig. 82. the shift in tilts of bedding in the piggyback basin of the br08 thrust sheet is interpreted to reflect the propagation of ramps. the strike is the same (110°), but the dip of the lower sand beds is 40°, corresponding to deposition during propagation along a flat, whereas the dip of beds above the truncation surface is only 28°, corresponding to deposition during ramping. photograph: june 1997. fig. 83. back-thrust reverse faults displacing the sand beds in the rubjerg knude formation deposited in the piggyback basin of the br08 thrust sheet. these reverse faults are interpreted to have formed during the thrust propagation of the footwall ramp of br06/br05. photograph: june 1997. 97 (fig. 77). in the brede rende section, this type of diapirism occurs commonly in the upper sand-rich part of the lønstrup klint formation and in the rubjerg knude formation. a good example occurs at the tip of the br05 thrust sheet, where the thrust fault (br04fwf/br05hwr) is completely obscured by hydrodynamic brecciation. examples of the second type of diapirism include the brede rende diapir and the kramrende diapir. here the clay-rich units of the lower part of the lønstrup klint formation became mobilised by over-pressured water (or gas) to form an intrusive grey, homogeneous mud. the diapirism in the brede rende section was formed syntectonically during ramping (pedersen 1987). the thrusting displaced some of the feeders in the mushroom-shaped diapirs, and some of the diapirs intruded through the thrust sheets up into the thrust sheet above. moreover, the mushroom-shaped diapirs penetrate the l/r-unconformity at the top of the diapir (fig. 73). it may also be noted that some of the diapir feeders have been tilted by the bending produced by ramp propagation (fig. 4; pedersen 1987). brede rende normal fault the brede rende normal fault (brhf), in the central part of the brede rende section, is a planar fault that strikes 100° and dips 45°s (fig. 75). the vertical displacement is c. 20 m when measured from the hanging-wall (thrust-fault) flat of the br06 in the footwall block of the normal fault to the same flat in the hanging-wall block of the brnf. a network of smaller normal faults with minor displacements occurs in the footwall block (fig. 80) and adds to the monoclinal bend in the footwall block of the brnf. in the hangingwall block, the br06 thrust sheet is dragged along the fault plane and the drag is bounded by a minor splay fault. moreover, a weakly developed rollover-anticline outlined by the br06 thrust sheet occurs in the hanging-wall block of the brnf (fig. 75). above the northern limb-bend of the rollover-anticline, a minor depression (c. 5 m deep) was formed. in this depression, a series of minor imbricate sandy mud slumps formed, which may be viewed as synto epitectonic deposits at the top of the rubjerg knude formation in br06 related to faulting of the brnf. frost wedges frost wedges or fossil ice wedges are recognised in the rubjerg knude formation, as preserved in the upper part of the br04 thrust sheet. the clif f section here became exposed after the cross-section (plate 1) was drafted and is thus not included. it would have been situated near point 3975 m in the cross-section. the frost-wedge fractures are 5–10 cm wide and are filled with structureless sand. along the sides of the fractures, the bedding in the sand is bent down towards the fracture due to minor displacements along small fanning normal faults; the vertical range of the frost wedges is about 1–3 m (fig. 81). the presence of frost wedges in the rubjerg knude formation clearly indicates that the sand was ground frozen, and thus also elevated above water level in the glaciofluvial and glaciolacustrine environment that prevailed during the deposition of the formation. the ground-frozen condition of the sand may be the reason for the excellent preservation of the normal fault network related to the brnf. interpretation of structural development the first thrust sheets to move were probably br03 and br06, which ramped up to the upper footwall flat and moved southwards over a thin cover of the rubjerg knude formation. in the balanced cross-section, the presence of the long, thin br03 thrust sheet, and especially br06, requires that there has to be underlying lower and intermediate duplex sheets. the balanced cross-section model favours a continuation in the subsurface of several segments of the lower thrust duplex. the br03 thrust sheet is viewed as a coherent thrust sheet, which from the ramp of the minor br02 thrust sheet, continues along the lower décollement surface at the 30 m level. the lower trailing-end duplex segment extends northwards to the thrust fault separating the brede rende and the sandrende sections (sr01hwr/br08fwr). this trailing segment of the br03 thrust sheet is estimated to be about 300 m long, and the remaining five thrust sheets in the brede rende section have all been ramped up onto this segment along which the allochthonous transport and piggyback displacement took place. the simplest model for understanding the framework of the duplexes is to accept segmentation of the trailing end of the br05 thrust sheet. it is necessary that br06 was thrust over br05 before the trailing 98 end of br05 was thrust up over the footwall ramp of br04. the existence of the intermediate br04 hanging-wall ramp indicates that br04 had to ramp up two footwall ramps in dif ferent positions of the trailing end of br03. thus the model indicates that a lower hanging-wall ramp of br04 was emplaced along the intermediate footwall flat of br03. according to the construction of the balanced cross-section, this also necessitates a lower duplex segment to be thrust up in front of the lower hanging-wall ramp and flat of br04. these dif ferential thrust displacements provide an explanation for the development of the brnf. the displacement along the brnf is consequently considered to be due to two factors. the first 10 m offset was caused by normal faulting in front of br05, where a foreland-dipping bend of the br06 hanging-wall flat was created over the nose of the br05 thrust sheet. the next 10 m displacement was caused by a foreland-dipping limb of the tip of a duplex segment situated beneath the br04 thrust sheet causing the br04 hanging-wall flat to act as a normal fault. at the north end of the section, the br07 thrust sheet, which only has 3–4 m of the rubjerg knude formation at the top, was overthrust by the br08 thrust sheet at an early stage. the thickness of about 10 m of rubjerg knude formation on top of br08 indicates that after the two thrust sheets were thrust-separated, deposition of rubjerg knude formation continued in the piggyback basin of br08. this sedimentation probably took place while br08 in a piggyback position on br07 ramped over a lower footwall ramp of br03 and propagated along an intermediate flat, passing over the footwall ramp of br05/br06, before the temporary cessation of thrusting. in the br08 piggyback basin, the propagation of ramps is reflected in the change in tilt of the bedding (fig. 82). moreover, a number of minor back-thrusts have been recognised in these beds (fig. 83), and are considered to have been related to the ramp propagation. in the dynamic development of the brede rende section, both the frontal southern and the northern parts were involved in diapirism. in both parts, it is fig. 84. the sandrende diapir developed in the sr02 thrust sheet. the arrow indicates the direction of reverse faulting, which marks the prominent back thrust. along the steep northern flank, the hanging-wall ramp of sr03 (sr03hwr ) was bent. the bend of the l/r-unconformity formed due to the fold-bend-folding of sr02 at the lower footwall ramp hinge. 99 evident that the diapirism was active after the emplacement of the thrust sheets, since the hanging-wall flats are penetrated by diapirs rising from a mobilised underlying thrust sheet. however, it is also evident that the diapirism ceased before the maximum compression of thrust sheets had occurred. the termination of thrust compression was reached when the maximum inclination of the flats occurred. this coincided with the conclusive accumulated ramping of piggyback thrust sheets. thus, the inclined position of the feeders to the mushroom-shaped diapirs indicates a synthrust intrusive emplacement. it is therefore concluded that the diapirism was activated by ramp propagation and that some of the diapirs can be regarded as extreme developments of hanging-wall anticlines created during soft sedimentary deformation (fig. 74). sandrende section the sandrende section is one of the most studied parts of the lønstrup klint cliff section (fig. 5; houmarknielsen et al. 1996; sadolin et al. 1997). even so, the development of this section is not fully understood, and some new and revised details are added here. the main feature of the section is a broad basin containing a thick succession of the rubjerg knude formation deposited in a piggyback basin. to the south, a diapir distorts this basin, and to the north the basin is over-thrust by a thrust sheet of the stenstue rende section. the central part of the section preserves a remarkable development of normal faults. these were formerly regarded to have formed in response to the volume adjustments in the sandrende diapir (sadolin et al. 1997), but are now interpreted as elements of a thrust-fault propagation model with differential duplex segments ramping in the subsurface. tectonic architecture the sandrende section comprises four thrust sheets (sr01–sr04). the southern boundary of the section is the trailing-edge ramp of br07 and br08 in the brede rende section, and the northern boundary is the rather steep (> 60°) trailing-edge ramp of sr04. the boundary with the stenstue rende section to the north is a combination of this trailing-edge ramp and the hanging-wall flat of the frontal southernmost thrust sheet in the stenstue rende section (see below). the transition between the brede rende section and the sandrende section in the subsurface is not clear due to uncertain relationships between br07– br08 and sr01. the description below is based on the preferred interpretation, which traces the trailingedge ramp of br08 in the brede rende section down to the décollement surface 30 m below the reference surface. this implies that the lowermost trailing ends of br07 and br03 remain as low-lying segments that sr01 had to ramp over. an extra segment and some smaller adjustment splints of the sr01 thrust sheet were also left in the subsurface. this is reflected in some of the structural features exposed in the section between sr01 and sr02. at the tip of the sr01 thrust sheet, the lønstrup klint formation forms a thin wedge, which indicates that the initial hanging-wall ramp (sr01hwr) only had a dip of about 10°. however, after ramping was concluded, the thrust fault was steepened to the present dip of 40°n; the measured orientation of the ramp is 108°/40°n. the displacement of sr01hwr along the upper footwall ramp and flat of br08 is c. 53 m. the frontal part of sr01 has a bend, and it only dips about 25°n due to the change in thrust-fault inclination passing the upper footwall ramp hinge and the subsequent introduction of a small satellite thrust fault displacing the lower part of sr01 up over the tipwedge. the consequence of thrusting the thin c. 50 m long frontal part of the thrust sheet is that in the balanced cross-section, a lower duplex segment (sr01u) must be accounted for, and that sr01u at an advanced stage of sr01 thrust propagation was picked up in the thrust translation (see below). the sr02 thrust sheet was formerly interpreted as a large-scale diapir (sadolin et al. 1997; fig. 84). in the present structural analysis, sr02 is treated as one large thrust sheet in which the mobilised mud underwent mud diapirism at a relatively late stage. this assumption permits an approximation of balancing the thrust sheets, accepting that the thrust faulting is evidently the most important part of the dynamic development. the argument for this is based on the fact that sr02 over-thrust the back of sr01 with a displacement of about 100 m. this 100 m of displacement has to be compensated for by the same amount of displacement along the lower décollement surface, which can be calculated to have taken place at a stratigraphic depth of 30 m below the l/r-unconformity. the thrusting of sr02 resulted in a considerable amount of elevation during propagation along footwall ramps (sr01fwr and br03fwr), since the l/r reference surface is sit100 uated about 35–40 m above sea level in the cliff section. thus the ramping and displacement along the upper flat took place before the final emplacement of the lower duplex segment of sr01, indicated by the normal fault displacement of both sr01 and the frontal part of sr02. as noted in the kramrende section description, steep ramping creates back-thrusting at the hinge of the hinterland-dipping limb. thus the peculiar mushroom-shaped structure with a wing pointing to the north is considered to be the effect of reverse faulting due to back-thrusting (fig. 84). the reverse fault feature may have been accentuated by re-orientated internal detachment folding and irregular diapirism in the sandrende diapir. furthermore, it should be noted that the l/r-unconformity surface has a steep dip on the northern flank of the sandrende diapir. near the beach level, the l/r-unconformity bends into a gentle dip indicating that in the trailing end of sr02, the lower hanging-wall flat rests on the lower footwall flat coinciding with the décollement level at 30 m. the sr03 thrust sheet is relatively small with a displacement of about 75 m. the tip of the thrust sheet is bent upwards into a nearly vertical position due to drag along the almost vertical northern flank of the sandrende diapir (sr02). thus the sr03 thrusting was rather early, but as the rubjerg knude formation is about 10 m thick in sr02 there was a significant time span before sr03 was thrust up on the back of sr02. sr03 was displaced up along the upper footwall ramp, which is exposed in the cliff section. sr03 was also displaced along an intermediate flat situated at the 20 m level, indicated by the thickness of the thrust wedge. during thrust propagation, the trailing end of sr03 was cut off and left as an isolated duplex segment, while the frontal part of sr03 was displaced along the intermediate flat (see plate 2). sr03 was over-thrust by sr04 with a relatively short time gap, as indicated by the thin (3 m) succession of rubjerg knude formation on top of sr03. the thrust displacement of sr04 over sr03 is about 60 m, and fig. 85. conjugate normal faults developed in the lønstrup klint formation in the sr04 thrust sheet. an offset of about 1 m can be recognised by correlating turbidite sand beds in the footwall block to the same beds in the hanging-wall block. the normal fault framework is interpreted to be due to lateral extension in the sr04 thrust sheet during its propagation over the upper footwall hinge of an underlying duplex. photograph: may 1995; measuring staf f divisions (centre) are 20 cm. 101 the accumulated displacement of the trailing end of sr04 relative to sr02 is in the order of 135 m. the frontal part of sr04 consists of a relatively thin wedge of the upper part of the lønstrup klint formation overlain by an up to 28 m thick succession of the rubjerg knude formation. in the central and rear parts of sr04, the thickness of the lønstrup klint formation increases to more than 20 m, indicating the existence of a hanging-wall ramp which can be traced down to the décollement zone, 30 m below the l/r reference surface. the central part of sr04 forms a broad hanging-wall anticline, where a number of extensional normal faults cross-cut the lønstrup klint formation (fig. 85). the southernmost normal fault in this system is considered to reflect the foreland-dipping features formed due to displacement of the hangingwall anticline along the intermediate flat. finally, it should be noted that the piggyback basin (rubjerg knude formation) of sr04 is divided into two subbasins, one in the southern frontal part and one in the northern trailing part of the thrust sheet. the area between the sub-basins lacks the rubjerg knude formation because it corresponds to the crest of the hanging-wall anticline. sedimentary units the type sections of the lønstrup klint and rubjerg knude formations, as defined in this bulletin and previously described by sadolin et al. (1997), are situated at sandrende. as defined above, this succession is divided here into the lønstrup klint formation and the overlying rubjerg knude formation, which are separated by the l/r-unconformity (fig. 19). the rubjerg knude formation is covered by an up to 1 m thick homogeneous sandy till, which is referred to the kattegat till formation (fig. 30). lønstrup klint formation in the sandrende section, the lower exposed part of the lønstrup klint formation is composed of laminated clayey to sandy mud, intercalated with a few thin sandy turbidites that grade up into finely laminated clay-rich mud. in the upper part of the formation, thicker turbidite sand beds with climbing ripples give the formation a banded light/dark coloured appearance (figs 19, 85). only very few load structures and hydrodynamic breccias have been noted in the sandrende section, except in the lower part of the sr04 sheet where ball-and-pillow structures and small-scale polydiapirism have been observed (fig. 86). the ball-andpillow features are about 20 cm in thickness, which is probably the thickness of the original beds; they are typically elongated about 50–75 cm parallel to the strike of the bedding, suggesting they were formed during the thrust deformation. rubjerg knude formation the rubjerg knude formation comprises three units: (1) a lower unit c. 5 m thick consisting of trough crossbedded sand and gravel, (2) a middle unit dominated by climbing ripple cross-laminated sand, and (3) an upper unit comprising alternating beds of small-scale ripple cross-laminated sand and trough cross-bedded sand (fig. 19). the units reflect the change from fluvial to lacustrine and back to fluvial depositional environments (sadolin et al. 1997). the rubjerg knude formation has an onlapping relationship in the frontal part of the sr04 thrust sheet, which reflects initial thrust faulting during sedimentation (sadolin et al.1997). in the central part of the sr04 thrust sheet, growth-fault sedimentation along normal faults is recorded in the lower part of the rubjerg knude formation. the growth faults coincide with the foreland-dipping limb of the hanging-wall anticline of sr04 (fig. 87). this syntectonic sedimentation supports the piggyback basin concept for deposition of the rubjerg knude formation. moreover, a slumped block 0.5 × 2 m in size occurs along one of the normal faults indicating that the tip of the satellite thrust in sr04 was exposed to erosion and slumped into the basin. similar slumped blocks were observed on the northern flank of the sandrende diapir indicating that the diapir rose above the depositional surface during emplacement and that fragments of the lønstrup klint formation slumped into the piggyback basin. the synsedimentary rise of the vertical diapir wall was also reflected in sedimentation of small point-bar wedges along the vertical flank of the sandrende diapir. towards the top of the rubjerg knude formation, broad trough cross-bedding is observed. minor thrust faults splaying out from the tip of ss01 displace the cross-bedded sand, and the base of some of the troughs dramatically truncate the thrust faults, in a similar fashion to that observed in the br04 piggyback basin of brede rende (see above). at the top of the rubjerg knude formation, sand was deposited in a 102 fig. 86. mobilisation and small-scale polydiapiric features developed in the upper part of the lønstrup klint formation in the sandrende section (trailing end of sr04). the polydiapirs started along a bed of clayey mud as small flames (with small wavelength), which were subsequently folded around the taller diapirs. the sandy beds above and below constitute planar laminated and climbing ripple cross-laminated fine-grained sand with organic debris and mud draping the ripples. locally in this sand, hydrodynamic mobilisation has created zones of mud-free structureless sand and the accumulation of mud forming dendritic structures. the dynamic development of the structure is illustrated in fig. 88. fig. 87. extensional normal faults with related growth-fault sedimentation of sand and gravel in the lower part of the rubjerg knude formation. the growth faults are marked with arrows indicating the direction of displacement. the top of the rubjerg knude formation in the sr04 thrust sheet is over thrusted by the ss01 thrust sheet. the two sheets are separated by a thrust fault that acts as footwall flat of sr04 (sr04fwf) and hanging-wall ramp of ss01 (ss01hwr). photograph: may 1995. 103 depression above the top of the sandrende diapir. this depression was probably formed by relaxation collapse of the diapir during consolidation and dehydration. structures and breccias structural investigations in the sandrende section focused mainly on the normal faults and their relationship to the thrusting, diapirism in the sandrende diapir, smallscale incipient polydiapirism, and the record of a deep frost wedge cutting the rubjerg knude formation. normal faults the sandrende section is an important locality for the investigation of normal faults formed on the forelanddipping limb of hanging-wall anticlines. thus, one set of normal faults displaces the frontal parts of sr01 and sr02, and another set displaces the central part of the sr04 thrust sheet. in the frontal part of sr01, a hanging-wall anticline developed due to ramping from the 10 m to the 20 m flat level. the normal faults here displace the rubjerg knude formation of sr01 as well as the tip of sr02. the faults now have a dip of about 45°s, but initially probably had a much steeper dip (up to 80°) subsequently reduced during the final bend of the br08 footwall flat. in the trailing end of sr01, a steep normal fault displaced sr01, as well as the over-thrust sr02, with an offset of 10 m. this probably reflects the influence of a sub-surface duplex, similar to the development of the brnf. the normal faults in the sr04 thrust sheet can be fig. 88. the polydiapiric structures and hydrodynamic breccias shown in fig. 86 are interpreted to have developed in the following five steps. 1: initial sedimentation of a clayey mud bed in a succession of mud and fine-grained sands. 2: first-order formation of small flames can be regarded as micro-diapirs with small wavelength. 3: second-order small diapirs developed with increased wavelength. small-scale thrusting and overturned geometry indicates formation during thrust-fault propagation. 4: increased mobilisation creates small-scale domes with extensional fractures forming in the crest. 5: liquefaction of the heterolithic sediment results in segregation of the sand and mud components. the mud accumulates in an irregular diapir from the top of which the mud-saturated liquid intrudes laterally along the primary parallel lamination. some mud and fragments of sand fall to the base of the diapir under gravity. 104 viewed as two sets of a fault framework. the first set formed 45–60°s dipping faults with displacements of 1–3 m. the southerly dipping tilt of the l/r-unconformity is regarded as the foreland-dipping limb of the hanging-wall anticline formed in sr04 (fig. 87). the faults above the foreland-dipping surface developed as growth faults associated with syntectonic sedimentation, as recorded in the lower part of the rubjerg knude formation. during displacement along the normal faults, a minor satellite thrust cross-cut the sr04 thrust sheet, and the tip of the satellite thrust sheet was slump-faulted to form slumped blocks in the growth-fault setting of the piggyback basin. the normal fault network at the crest of sr04 is very impressive (fig. 85). the strike of the normal faults is 090° with a dominant dip of 50°s, although a small number of conjugate faults with a dip of 60– 75°n also occur. in view of the angle of conjugate faulting, these normal faults could have formed due to the loading of the ss01 thrust sheet emplaced above the upper footwall flat of sr04, but could also have formed due to necessary extensional adjustments during propagation over the hinge of the footwall ramp. diapir structures the sandrende diapir only af fected one thrust sheet (sr02), in contrast to the brede rende and kramrende diapirs where two or more thrust sheets were involved in the diapir formation. the most impressive feature of the sandrende diapir is the major back-thrust, which has an offset of about 20 m towards the north. initially it was probably an almost vertical reverse fault, which was re-orientated and accentuated during thrustfault propagation. a number of smaller reverse faults occur along the steep northern wall of the diapir, which internally is composed of mobilised mud. in the upper part of the diapir, distorted bedding-structures isolated as ‘xenoliths’ in the upper part of the diapir are interpreted as fragments of hanging-wall anticlines. judging from the thickness of the diapir feeder, the diapir formed over a hanging-wall ramp where the sr02 thrusting ramped from the upper 10 m flat to the lower 20–30 m flat level. the formation of diapirs was evidently initiated by mobilisation on a small scale. an illustrative small-scale example of diapirism was observed in the northern part of the sr04 thrust sheet (fig. 86) where mobilisation and small-scale polydiapirs developed in the upper part of the lønstrup klint formation in the sandrende section (trailing end of sr04). the polydiapirs are related to beds of clayey mud deposited between the thicker beds of sandy turbidites. along the boundary of the 25–75 cm high diapirs, small flame structures occur and the laminated sandy beds above are irregularly folded. locally, hydrodynamic mobilisation created mud-free structureless sand and complex mud structures developed. an interpretation of the dynamic development of the structures is given in fig. 88. frost wedge a 20 m deep fracture cross-cuts the rubjerg knude formation in the central part of the southern sub-basin in the sr04 thrust sheet. the fracture is less than 10 cm across, and can be followed as an irregular trace downwards into the sand sequence with a number of minor lateral jumps. this irregular fracture is one of the few structures that can be interpreted as a frost wedge. it does not penetrate the overlying ss01 thrust sheet, indicating that it formed within the rubjerg knude formation from an exposed surface downwards into a freshly frozen sand package. it can be inferred that during the latest phase of thrusting, the sr04 thrust sheet was elevated to a position such that the top of the rubjerg knude formation was exposed subaerially. interpretation of structural development a hanging-wall anticline developed c. 40 m from the tip of sr01 when it passed the footwall ramp and flat of br08. this initially created a foreland-dipping tilt of the sr01 thrust structures, and was also responsible for the formation of the normal faults described above. however, the frontal part of sr01 has to be accommodated with a duplex segment in the subsurface (sr01u). the trailing end of sr01 is rooted down to the décollement level, where it corresponds to the segment adjusting the c. 80 m long frontal part of sr02. when the hanging-wall ramp of sr02 initiated the propagation up along the footwall ramp, a hangingwall anticline was formed that developed into the sandrende diapir with its marked back-thrust. during the sr02 propagation along the footwall ramp, the sr01u-duplex was pushed up and created a minor hanging-wall anticline, along which foreland-dipping limb a normal fault developed and displaced the sr02 thrust sheet as well as sediments in the piggyback basin of sr01. 105 the lower 10 m of the rubjerg knude formation was deposited throughout the sandrende section, with the exception of the northern part of sr03, which had already been blocked by thrusting of sr04. displacement of the lower hanging-wall ramp of sr04 onto the intermediate flat of sr03 (and sr02) then took place. subsequently, the first normal growth faulting was initiated at the margin of the southern part of the piggyback basin above the foreland-dipping l/r-unconformity. propagation of the lower sr04 hanging-wall ramp separated the piggyback basin into two sub-basins where deposition of the upper part of the rubjerg knude formation took place, while the crest of the anticline between the sub-basins was probably subjected to erosion. the southernmost thrust of the stenstue rende section (ss01) over-thrust the top surface of the rubjerg knude formation (sr04fwf) as well as the eroded surface of the ramp anticline; this prevented deposition in the sr04 piggyback basin. ramping of the lower trailing segment of sr03 was activated in the latest stage of thrusting. the propagation of this duplex segment (sr03u) for a short distance up along the footwall ramp contributed to the flat-topped hanging-wall anticline formed in sr04. this final duplex emplacement may have been one of the causes for the formation of the normal fault framework in the hanging-wall anticline of sr04 (fig. 85). stenstue rende section in the stenstue rende section, a remarkable and dramatic episode of megaslumping is recorded. formation of a very large southward-verging anticline was accompanied by chaotic hydrodynamic brecciation (fig. 89). another important element related to this section is the c. 200 m displacement of the frontal thrust sheet over the sandrende section to the south. fig. 89. the large slump fold in the stenstue rende section. the slumping folded the ss05 thrust sheet into an overturned anticline during displacement down the normal fault escarpment. the escarpment was formed during normal fault displacement of the tip of ss04 parallel to the foreland-dipping limb of a hanging-wall anticline in ss03 (see plate 2 and fig. 90). photograph: june 1999. 106 the stenstue rende section is named after the gully situated between the stenstue rende section and the sandrende section leading inland from the beach. in the northern part of the section is the gully known as the søndre grønne rende. this is reached by a path through the pinewood connecting with the main road between rubjerg and lønstrup. tectonic architecture the stenstue rende section comprises six thrust sheets (ss01–ss06). to the south, the footwall ramp and flat of sr04 in the sandrende section bound the section. to the north, the boundary is defined by the footwall ramp of ss06, which coincides with the hanging-wall flat of the southernmost thrust in the grønne rende section (gr01). the most important thrust sheet in the stenstue rende section is the more than 400 m long ss01 thrust sheet, the frontal part of which over-thrust the northern part of the sandrende section and has a displacement of more than 200 m. the initial ramping of the ss01 thrust sheet was located at a gently dipping hanging-wall ramp. subsequent to the ramping, part of the tip was eroded away during the uplift exposure of the hanging-wall anticline above the ramp and the final truncation of the glaciotectonic unconformity. due to the increase in thickness of the ss01 thrust sheet, corresponding to a change from the 10 m upper flat level to the 20 m flat level, an intermediate hanging-wall ramp developed about 100 m from the frontal tip. this hanging-wall ramp rests on top of the footwall flat (sr04fwf) above the prominent normal fault structure in the sandrende section. only a small remnant of the northern part of the upper flat structure is preserved, and this is not very well exposed due to its location in the inner part of the stenstue rende. the lower ss01 hanging-wall ramp (ss01hwr, ramping from the 30 to 20 m flat level) situated in the middle part of the ss01 thrust sheet is now exposed in a steeply dipping position along the sr04 footwall ramp. the propagation of ss01hwr was responsible for the bend of the footwall syncline in the rubjerg knude formation in sr04, and the subsequent tilting of ss01hwr resulted in the appearance of a more or less vertical boundary between the two sections. the vertical orientation is a combination of 45° dip on the footwall ramp added to 45° dip on the hanging-wall ramp. note that in the balanced cross-section, there is a c. 200 m long lower ss01 duplex segment (ss01u) which needs to be allowed for. this implies that after ramp propagation, the lower hanging-wall flat of ss01 was displaced along the intermediate footwall flat on ss01u. ss02 is a small thrust sheet, thrust onto the footwall ramp of ss01; this footwall is composed of the rubjerg knude formation situated in the upper part of ss01. note that the frontal part of ss02 has a surprising vertical orientation and is displaced by a more or less horizontal extensional fault, the cause of which is discussed below. the frontal part of the ss03 thrust sheet is shown in the cross-section as a rather simple, upright thrust structure (plate 1). however, the ss03 thrust sheet is in reality a chaotic load and hydrodynamic breccia complex. at the base of the cliff section is an upright anticline, which is regarded as a key structure in the interpretation of the thrust development (fig. 90). above the anticline, a normal fault dipping 40°s truncates the c. 30 m thick rubjerg knude formation. the thin frontal part of the ss04 thrust sheet is characterised by chaotic brecciation. the trailing end is c. 10 m thick, dipping 45°n, with the hanging-wall flat thrust along the footwall flat of ss03. the tip is separated from the trailing part of the ss04 thrust sheet by a 40° dipping normal fault with a displacement of about 50 m. this normal fault formed an escarpment truncating the rubjerg knude formation on top of the ss03 thrust sheet, upon which deposition of a coarse clastic breccia took place (fig. 91). the normal fault escarpment was finally overridden by the frontal part of the ss05 thrust sheet, which slump-thrusted down the fault plane and formed a major overturned slump fold (fig. 89). the formation of the megaslump fold took place after the ss05 thrust sheet was thrust up along the footwall flat of ss04 to the head of the escarpment from where it gravitationally slid down to the depression on the back of the ss04 tip. a soft sedimentary tectonic breccia was formed at the transition between ss04 and ss05, which was cross-cut by a number of minor steeply southward dipping normal faults reflecting the final settling of the fault-slump structure. the ss06 thrust sheet is about 30 m thick, its lower hanging-wall flat resting on the footwall ramp of ss05. it has a steep dip and has been strongly disturbed by mobilisation and internal diapirism. this thrust sheet is included in the stenstue rende section because it involves the trailing lower duplex segments of ss05 and ss04. from the position of the l/r-unconformity surface, about 30 m above sea level, it can be inferred that the ss06 thrust sheet was raised up over the low107 er trailing segments during ramping and subsequent stacking of a subsurface duplex complex. sedimentary units the most interesting sedimentological feature within the stenstue rende section is the record of syntectonic sedimentation related to normal faulting. this includes slump deposits as well as a gravel bed developed on the escarpment surface of the normal fault. as these sedimentary features are related to deposition in the piggyback basin, they are described below as part of the rubjerg knude formation. the sediments of the lower lønstrup klint formation have been strongly affected by thrust shearing, and the upper levels were modified by hydrodynamic brecciation. the rubjerg knude formation comprises a confusing mixture of redeposited units together with the main fluvial-lacustrine sediments related to the piggyback basins. lønstrup klint formation the lower part of the lønstrup klint formation is exposed in the ss06 thrust sheet, where the lower hanging-wall flat is thrust up along the footwall ramp of ss05. here bluish grey clay alternates with dark redbrown clay in a laminated to thin-bedded unit (fig. fig. 90. the crest of the hanging-wall anticline formed in the ss03 thrust sheet in the stenstue rende section. photograph: june 1984; the staf f divisions are 20 cm. fig. 91. the conglomerate/breccia formed along the fault escarpment truncating the ss03 thrust sheet. photograph: june 1984; the staff divisions are 20 cm. 108 92). it is evident from the shear structures that this unit acted as a décollement zone during thrusting. the main part of the lønstrup klint formation exposed in this section comprises the upper sand-dominated part of the succession. the breccias in the ss03 and ss04 thrust sheets probably initially formed as mediumto large-scale ball-and-pillow structures in sand beds 20–60 cm thick during initial thrusting; the formation was subsequently deformed during gravity slumping. rubjerg knude formation nearly 20 m of fluvial-lacustrine sand were deposited in the piggyback basin of ss01 and ss03 during the thrust-fault activity af fecting the stenstue rende section. however, the most conspicuous unit is the remarkable conglomerate/breccia related to the normal fault. the gravel bed draping the fault escarpment is 10–50 cm thick and includes clasts up to 10 cm in size. locally, the clasts occur in a clayey mud matrix, but the latter has often been removed by recent erosion. it is perhaps surprising that a coarse gravel bed could have accumulated and been preserved along a fault escarpment dipping at about 35° (fig. 91). one explanation may be that the escarpment was only exposed for a very short time before the ss04 thrust sheet was displaced down the fault plane; in this case, the redeposited gravel rather represents a tectonic breccia composed of the smeared-out lithologies of the l/r-unconformity and surrounding sediments. the breccia is thus interpreted as the residue of a brecciated thrust sheet. the source of the clasts was probably the l/r-unconformity, and some of the material may have been derived from the unconformity by successive erosion during exposure at the head of the fault escarpment. this probably only occurred for a brief period before the escarpment was covered by the slump-slide of the ss05 thrust sheet. the small piggyback basin on top of the ss05 thrust sheet is a double syntectonic basin which was partly carried on the back of a thrust sheet as well as acting as a depression in the hanging wall of a normal fault. a 9 m thick succession represents the fill of this basin. the lowermost 3 m consist of large-scale cross-bedded medium-grained sand, rich in clay and silty mud clasts. towards the upper part of this unit, clay drapes on the cross-bed foresets become more common and the beds are affected by small-scale slumping. the overlying 5 m thick unit comprises sand beds 30–50 cm thick, with mud intercalations 5–20 cm in thickness. clay clasts are common and the sand shows smallscale ripples. the uppermost 1 m thick bed consists of mainly horizontal laminated sand and mud. this piggyback basin succession is interpreted to record a fluvial depositional environment that with time developed into a small shallow lake. a number of small south-dipping normal faults intersect the rubjerg knude formation up to the base of the thinly bedded muds and sands, indicating that the lake first became established when the fault activity ceased. structures four types of structures in the stenstue rende section deserve particular mention: (1) mesoscopic thrust-fault structures above the lower hanging-wall flat, (2) hanging-wall anticlines, notably the one in the central part of the section, (3) normal faults, the most important being the major escarpment-producing fault, and (4) slump folding related to the escarpment of the same fault. thrust-fault structures thrust faulting related to the décollement zone in the stenstue rende section has been observed in the lower hanging-wall flat of the ss06 thrust sheet. here the décollement zone is located in the 30 m flat level, which corresponds to the base of the 30 m thick lønstrup klint formation where lithologies are mud-dominated, comparing dark blue-green-greyish, clayey or silty mud with a few light grey coloured, fine-grained sand laminae. two types of structures are distinguished: imbricate duplexes and listric imbricate fans (figs 92, 93). the duplex imbricates appear within a 1 m thick unit bounded by thrust-shear surfaces below and above (fig. 92). the mesoscopic-scale duplex complex consists of sheets about 0.5 m thick and 1–3 m long. some of the duplexes are folded into antiformal stacks and form lensoid networks. the basal and roofing thrust faults occur as 20 cm thick shear bands penetrated by flat anastomosing jointing (fig. 92). the listric fans are outlined by 1–2 cm thick sedimentary layers or tectonically induced sand streaks (fig. 93). they rise from a narrow thrust plane, recognisable as a joint surface draped by a 1 mm thick film of black mud, and extend upwards into the muddy lithology where they seem to disappear before being over-thrust 109 by the next thrust joint surface about 1 m above the basal thrust surface. hanging-wall anticlines three hanging-wall anticlines have been recognised; the anticline in the frontal part of ss01 has been commented on above. the second example is not very obvious, but was developed above the intermediate hanging-wall ramp of ss01. the structures related to it were later modified by re-orientation due to the bend of ss01 up along the sr04 footwall ramp. the third hanging-wall anticline is the key structure in the stenstue rende section and is situated in the middle part of the ss03 thrust sheet. the ss03 hanging-wall anticline (fig. 90) was folded due to the ramping in the middle of the lower trailing duplex segment (ss01u). this ramping took place at a mature stage of thrusting, and ss01u was separated into two segments. the anticline is upright and tight, and the onlapping sedimentation of the rubjerg knude formation on the northern flank indicates that the ss03 thrust sheet had commenced transport along a footwall ramp and flat prior to the anticlinal folding. fig. 92. along the lower hanging-wall flat of the ss06 thrust sheet, an imbricate duplex complex has been recognised; bounding thrusts indicated by shear arrows . the thrust-fault imbrication formed in the lowermost part of the lønstrup klint formation during displacement along the décollement surface. the trowel is c. 30 cm long. photograph: june 1997. fig. 93. an imbricate fan formed in the lower part of the lønstrup klint formation in the ss06 thrust sheet. the trowel is c. 30 cm long. photograph: june 1997. 110 normal faults two normal faults are discussed: (1) the extensional fault with horizontal fault plane that displaces the tip of ss02, and (2) the major normal fault displacing the ss04 thrust sheet. the extensional fault affecting ss02 was formed north of the hanging-wall anticline developed over the intermediate hanging-wall ramp of ss01. it is interpreted to have formed initially as a normal fault dipping c. 45°n on the foreland-dipping limb of the ss01 hanging-wall anticline. subsequent to displacement on the normal fault, the ss02 thrust sheet and the fault were tilted into vertical and horizontal positions, respectively, during the fault-bend folding resulting from the ss01 propagation up along the footwall ramp. the major normal fault displacing ss04 is also regarded as a fault that developed on the foreland-dipping limb, here related to the anticline in ss03. it is observed that the l/r-unconformity dips beneath the beach level, indicating that the reference surface is not elevated and consequently that the underlying ss03 hanging-wall flat rests on a footwall flat; the normal fault is thus preserved with its initial orientation. the formation of the normal fault is similar to the formation of the brnf in the brede rende section (see above). slump folding the large-scale slump fold formed by the ss05 thrust sheet as it was displaced down the foreland-dipping fault escarpment can be compared to the same type of deformation described from the martørv bakker section. however, in the stenstue rende section, the lønstrup klint formation is still preserved as a coherent sheet, deformed into a major southerly overturned fold with an amplitude of about 15 m and an irregular fold axis orientated se–nw (c. 150°) (fig. 89). during slumping along the escarpment, the redeposited units were strongly affected by hydrodynamic brecciation resulting in the chaotic disorganised nature of the sediments. interpretation of structural development the important question in the development of the stenstue rende section is the time of formation of the ramp bend anticline during thrust-fault propagation. the interpretation given here is based on the description above, and the balanced cross-section and model for ramping in the subsurface given in the ramp cross-section (plate 2b). firstly, it should be remembered that there is evidence of a long translation along the décollement zone, primarily indicated by the considerable distance of ss01 transport over the upper footwall flat in the sandrende section (sr04). secondly, the displacement of the frontal part of ss01 must be balanced with a lower duplex segment (ss01u) in the subsurface. moreover, the displacement of ss01 also affected the ss02 thrust sheet by superimposed structural development. the superimposed model here advocated is supported by the following interpretation. as the initial angle of thrust faulting rarely exceeds 30° (jaeger & cook 1979), superimposed rotation must have affected the ss02 thrust. the angle between the ss02 hanging-wall ramp and the l/r-unconformity surface is about 30° indicating a normal type of thrusting when the rubjerg knude formation was horizontal. considering the thrust in this pre-rotated position, it is easy to envisage that the extensional fault offsetting the tip of ss02 as a normal fault formed over the lower hanging-wall ramp of ss01. to restore the thrust sheet into an upright position, two phases of rotation are necessary. the first one would be the ss01 ramping on the footwall ramp of sr04, and the second would be the re-orientation of the ramp due to the fault-bend provided by the thrusting of a subsurface segment of sr03 up along its footwall ramp in the sandrende section. the initial ramping of the leading edge of ss01 probably took place during sedimentation in the lower part of the piggyback basin of sr04. thrust propagation of ss02 must have been initiated at the same time, indicating that the ss03 thrust sheet in the trailing end of ss02 also participated in the translation along the 10 m flat level (ss01 intermediate footwall flat). during the translation of the lower footwall ramp in the trailing end of ss01, the 200 m long lower ss01u segment must also have been thrust, which is interpreted to have caused the ramping in the central part of ss01u. above this ramp, a lower hanging-wall anticline developed, which also folded the overlying ss03 thrust sheet into the exposed anticline in the middle part of ss03, resulting in the foreland-dipping footwall flat of ss03 and the initiation of normal faulting. part of the ss04 thrust sheet had by then already propagated over the ss03 footwall flat, and was therefore subsequently displaced by the normal fault with a drag down the fault plane. the displacement on the ss01u footwall ramp must have been relatively small to create and preserve an upright, close to tight anticline. if the dis111 placement had continued, it is likely that a more flattopped anticline would have developed and minor normal fault imbricates would have been the result, rather than the marked fault escarpment that actually formed on the southern flank of the anticline. the slump-thrusting of ss05 down into the depression on the hanging-wall block of the normal fault is interpreted to have taken place shortly after the ss04 was down-faulted. propagation of the ss05 thrust sheet was combined with the push on its footwall ramp by thrusting of the ss06 sheet. during this final thrusting in the stenstue rende section, the trailing lower segments were stacked into a duplex, probably analogous to the mesoscopic-scale duplex structure exposed along the hanging-wall flat of ss06 (fig. 92). grønne rende section in the geological cross-section of lønstrup klint presented by jessen (1931), two gullies were indicated south of the rubjerg knude fyr (the lighthouse), namely søndre and nørre grønne rende. by the year 2000, the cliff profile had no obvious gullies that these names can be attached to, although søndre grønne rende must have been close to the gully so annotated in the northern part of the stenstue rende section. the name grønne rende section is therefore adopted here to cover the section between stenstue rende and rubjerg knude fyr. the section comprises twelve nearly vertically orientated thrust sheets, which can be characterised as a listric imbricate fan. only the frontal parts of the thrust sheets are ramped up into steeply dipping positions, and in these parts of the thrust sheets the lønstrup klint formation is thin, whereas the rubjerg knude formation is relatively thick. the general impression of the section is of thin mud sheets alternating with thick units of sand (fig. 94). the points of interest in this section are the means of formation of an imbricate fan of uniform thrust sheets, and the mechanism by which the sheets reached their vertical orientation. also of interest is the arrangement of the now concealed duplex segments in the subsurface, where balancing of the thrust sheets indicates shortening of about 60%. fig. 94. view along the grønne rende section to the north where the lønstrup klint formation forms thin mud sheets interleaved with thick sand sheets referred to the rubjerg knude formation. in the far distance, stortorn forms the vertical cliff facing the sea. height of cliff is c. 50 m. photograph: august 1984. 112 tectonic architecture the grønne rende section comprises a leading-edge thrust sheet (gr01), which consists of a 30 m thick section of the lønstrup klint formation, succeeded above the l/r-unconformity by about 15 m of the rubjerg knude formation. north of gr01, a further twelve thrust sheets (annotated gr02–gr13) are exposed, each composed of an average thickness of c. 10 m of the lønstrup klint formation overlain by about 25 m of the rubjerg knude formation. to the south, the section is bounded by the thrust fault that separates the lower hanging-wall flat of the grønne rende frontal thrust sheet (gr01) from the footwall rampand-flat of the northernmost thrust sheet (ss06) in the stenstue rende section. to the north, the boundary of the grønne rende section is defined by the thrust fault that acts both as the hanging-wall ramp of rf01, the frontal thrust sheet in the rubjerg fyr section, and as the footwall ramp-and-flat of gr13. in the description of the thrust sheets, it is assumed that the thrust sheets initially involved only the lønstrup klint formation, and that the rubjerg knude formation was deposited syntectonically and separated into small piggyback basins between the sheets. as the imbricate thrust sheets constitute the most important structural element in this section, each thrust sheet is described separately in the structural account below. sedimentary units in the grønne rende section, the lønstrup klint formation is mainly represented by the upper levels of the formation. the only exception is the southernmost thrust sheet gr01, in which lower stratigraphic levels of the formation are also exposed. in this section, intra-rubjerg knude formation erosional surfaces locally incise the unconformity defining the lønstrup klint formation – rubjerg knude formation boundary. this unconformity is thus composite in places but the term l/r-unconformity is retained as it clearly still forms the boundary between these two formations. during hanging-wall ramping of the thrust sheet tips, the gravel beds on the unconformity were partly removed from the unconformity surface, and desiccation cracks may be present (only observed in the uppermost tips of the thrust sheets) indicating that some of the tips were exposed above water level. the rubjerg knude formation mainly comprises the same three units described in the sandrende section. however, a number of variations in sedimentary architecture occur due to syntectonic sedimentation. lønstrup klint formation the lower part of the formation exposed along the hanging-wall flat of gr01 consists of dark grey laminated mud with a few c. 0.5 m thick white sand turbidites; these have been strongly disturbed by thrusting, contortion and mud-mobilisation. the upper part of the formation is dominated by light-coloured sandy turbidites up to 1 m thick, interbedded with 10 cm layers of blue-grey clayey mud. rubjerg knude formation it has already been noted that the rubjerg knude formation was deposited in a number of small piggyback sub-basins. in the description of the piggyback basin architecture, four depositional elements are differentiated. 1. flat-parallel bedding (f-bedding): initially horizontal stratification of a bed deposited on a surface parallel to a flat as well as to the mean level of the l/r-unconformity. 2. ramp onlap (r-onlap): horizontal stratification or large-scale cross-bedding in a bed deposited on an inclined unconformity surface that had been tilted due to ramping prior to sedimentation. 3. foreland-dipping onlap (d-onlap): initially horizontal stratification in a bed deposited on an inclined unconformity surface dipping towards the foreland due to the repositioning of a hangingwall ramp on a footwall flat. 4. climbing ripple stratification (c-bedding): climbing ripple cross-lamination in beds 20–80 cm thick, commonly limited by f-bedding below and above (figs 24, 95). the rubjerg knude formation is interpreted as a glaciolacustrine deposit. sediment influx was probably relatively constant, and the sedimentary structures developed in the individual thrust sheets were governed by local conditions. during thrusting, the sedimentary base level changed, and the accommodation space varied depending on the size of the piggyback basins. thus the flow regime fluctuated and a variety of sed113 imentary structures formed, which are interpreted to reflect the thrust-fault development. the syntectonic banana-shaped basin described by pedersen (1987; fig. 4) was based on observations in the upper part of these piggyback sub-basins. these structures might also be characterised as footwall synclines that developed as growth-fault synclines, where deposition took place as the hanging-wall block was thrusted up along the footwall ramp dragging the underlying limb up along the thrust fault during the displacement. structures a systematic description of each thrust sheet in the section is provided below, together with some references to the syntectonic sedimentation. in general, the thrust sheets constitute an upright hanging-wall ramp, which initially had a dip of less than 20°. hydrodynamic brecciation and mud mobilisation occurred along the hanging-wall ramps and flats. along the upper footwall ramp, footwall synclines with compressive deformation of climbing ripple cross-laminated sands are very common, mainly developed as growthfault synclines as mentioned above. all the thrust sheets from gr02 to gr13 can be demonstrated to have been carried piggyback on the gr01 thrust sheet. gr01 thrust sheet the accumulated displacement of gr01 is estimated to about 140 m. this includes an interpreted displacement, c. 40 m, of the thrust-sheet tip. the thrust-sheet tip was, at a post-thrust stage, eroded away by the truncation of the glaciotectonic unconformity; calculation of the displacement of the tip follows the principle illustrated in fig. 11. the total displacement also includes the c. 100 m displacement along the exposed footwall ramp of ss06 and its consequent continuation down to the décollement zone in the 30 m flat level (stratigraphic level from the l/r-unconformity). from the base of the cliff and down into the subsurface, thrusting took place along the gr01 hangingwall flat. the internal tectonic structure of the lønstrup klint formation in gr01 is very similar to the thrust structures described from the ss06 thrust sheet of the stenstue rende section (figs 92, 93) with intraformational duplex structures, imbrication and strong mobilisation along the hanging-wall flat. as can be seen from the cross-section (plate 2b), the l/r-unconformity is traceable down to a level c. 5 m below sea level. this interpretation is supported by field observations, although the base of the clif f is often scree covered, and implies that the hanging-wall flat can be traced down to the décollement surface, and that the thrust sheet has not been elevated up onto, and translated along, intermediate flats in the subsurface. the l/runconformity rests in its initial stratigraphic position, and the reference level lies below sea level. the accufig. 95. large-scale cross-bedding displaying d-onlap overlain by planar bedding (d-onlap) and climbing ripple cross-laminated sand (c-bedding) of the rubjerg knude formation in the grønne rende section. photograph: july 1999; way-up is to the left. l/r-u, l/runconformity. 114 mulated thickness of the rubjerg knude formation is c. 15 m, and the sand beds were deposited with an onlap onto the northerly dipping l/r-unconformity (r-onlap). gr02 thrust sheet the gr02 thrust sheet is the southernmost imbricate in the imbricate fan of the grønne rende section. it consists of a 10 m thick section of the lønstrup klint formation overlain by an about 30 m thick section of the rubjerg knude formation. the main part of the thrust dips at 55°n, whereas the upper part is somewhat steeper. the lower 5–8 m thick unit of the rubjerg knude formation is characterised by large-scale trough crossbedding, and r-onlap can be recognised. f-dipping bedding, grading up into d-onlap in the uppermost part of the clif f section overlies this lower unit; this indicates ramp–flat–foreland dipping relationships during ramp and flat propagation. above this, a middle sand unit with f-bedding was deposited, and finally the upper unit shows r-onlap, which was subsequently folded in a footwall syncline below the gr03 thrust fault. gr03 thrust sheet in the gr03 thrust sheet, the lower part comprising the lønstrup klint formation is a wedge-shaped structure aligned along a vertical thrust fault (gr03 hanging-wall ramp). the thickness varies from c. 15 m in the lower part to c. 3 m in the upper part. above the l/r-unconformity, the rubjerg knude formation consists of a more than 30 m thick succession, which indicates that the piggyback basin of gr03 (as well as gr02) was a long-lived depocentre. the geometry of the gr03 hanging-wall ramp implies that it can be traced down to the 15 m (or 20 m) level. during ramping, a second stage of erosion affected the l/r-unconformity, which thinned out the lønstrup klint formation (fig. 96). the r-onlap in the lower half of the rubjerg knude formation probably reflects the ramping on the footwall ramp of gr02, and the middle part of the rubjerg knude formation was deposited during the propagation of the hanging-wall ramp over the footwall flat of gr02. the uppermost 4 m of the gr03 thrust sheet is very disturbed, probably due to push from the gr04 upper hanging-wall ramp. gr04 thrust sheet in the gr04 thrust sheet, the lønstrup klint formation is relatively thick, c. 20 m in the lower part of the cliff section and about 7 m in the top part. the gr04 hanging-wall ramp dips 70–80°n, and in the middle part of the cliff section a small hanging-wall ramp about 5 m high is preserved. in front of this ramp, the sand deposited at the top of the gr03 piggyback basin was pushed forward during thrust faulting along the upper footwall flat, as mentioned above. fig. 96. the composite development of the l/r-unconformity (l/r-u) resulted in the reduction of the thickness of the gr03 thrust sheet to a very thin horizon interlayered with thick piles of sand referred to the rubjerg knude formation in the grønne rende section. the upper hanging-wall ramp (gr03hwr) displays marked relief due to erosion. photograph: october 2000. 115 the maximum thickness of the rubjerg knude formation in gr04 is similar to the thickness in gr02 and gr03, but the piggyback basin is wedge-shaped due to the footwall ramp produced by the thrusting of gr05. the large-scale trough cross-bedded lower part of the formation tends to show r-onlap towards the upper part of the l/r-unconformity. gr05 thrust sheet gr04 and gr05 initially formed one coherent thrust sheet, with gr05 being carried piggyback during thrust propagation of gr04 before they were separated by the satellite thrusting of gr05 up along the footwall ramp of gr04. the hanging-wall ramp drops from the 5 to the 10 m flat level along a relatively steep ramp, which lifted gr05 free of gr04. in the lower part of the cliff section the l/r unconformity is situated about 5 m above sea level, indicating ramping to an intermediate flat in the subsurface. the gr05 thrust fault is overturned to the north, indicating that the thrust plane dips at 75°s (fig. 97). the uppermost 10–15 m of sand beds in the gr05 piggyback basin are horizontally orientated. the sand beds show large-scale trough cross-bedding and sedimentation is inferred to have taken place between the gr04 and gr05 thrust tips when these were exposed above the sediment/water interface during the latest stage of dynamic development. gr06 thrust sheet the lønstrup klint formation of gr06 is generally a relatively thick unit (10–15 m) although locally in this section deep erosion is evident at the l/r-unconformity. this localised deep erosion at the unconformity was probably due to erosion of a hanging-wall anticline during propagation over the footwall ramp hinge. this suggestion is supported by the presence of r-onlap in the lower to middle part of the cliff. the cfig. 97. the thrust-fault structures related to the gr04 and gr05 thrust sheets. the ‘overturned’ orientation of the gr05 hangingwall ramp (gr05hwr) is regarded as the result of repeated footwall ramping of the gr04 thrust sheet (gr04fwr = footwall ramp of the gr04 thrust sheet), and subsequent translation of the imbricate fan along the lower décollement surface. photograph: july 1999. 116 bedding observed in the middle part of the rubjerg knude formation may well reflect post-ramp deposition; subsequently, sedimentation briefly took place during displacement on the upper footwall flat. the thickness of the rubjerg knude formation is 15 m where the unconformity is deeply incised, decreasing to only 10 m laterally. this indicates that the basin was partly closed by the gr07 thrust propagating along its hanging-wall flat (back of gr06) in an early phase of development of the section. the uppermost beds show a high-angle r-onlap to the unconformity in gr06, demonstrating that the sand was deposited during the final phase of upthrusting, and just before the last c. 30° tilting of the thrust sheet into its present upright position. gr07 thrust sheet the gr07 thrust sheet consists of a 10 m thick unit of the upper part of the lønstrup klint formation with medium-bedded light grey fine-grained sand interbedded with thin mud layers. the l/r-unconformity is parallel with the hanging-wall flat in the main part of the exposed thrust sheet giving the impression that the sheet is of uniform thickness. the topmost part of the sheet is wedge-shaped where the upper hangingwall ramp is preserved, and the irregular structures of the tip can be interpreted as an upper hanging-wall anticline. the rubjerg knude formation in gr07 is about 15 m thick and can be divided into three 5 m thick units, which show the typical characteristics of sedimentation in the formation. in the uppermost part of the piggyback basin, high-angle r-onlap, similar to the bedding in the gr04–gr06 thrust sheets, indicates late syntectonic deposition between the thrust-sheet tips. gr08 thrust sheet the lower part of gr08, comprising the lønstrup klint formation, forms a wedge-shaped structure, with a thickness of only 3 m in the top of the clif f section and about 10 m at the base. the basal part is characterised by mobilised mud bounded at the thrust sole by a vertical thrust fault. at the top, a hanging-wall anticline and small diapir deformed the tip. the rubjerg knude formation of the thrust sheet comprises three units. the lowermost unit, up to 10 m thick, shows r-onlap in the lower part which is also the lower part of the cliff section. upwards, along the steeply dipping l/r-unconformity, the dip of the ronlap increases. there is an angular discordance between these beds and the beds occurring above. these beds show planar parallel bedding (f-bedding). in the lower part of this f-bedded unit, clasts of mud occur with sizes from cobbles to boulders (1 m size). these boulder-sized mud-blocks are interpreted as fragments of the gr09 thrust tip that were deposited by gravity slumping in the piggyback basin. the middle unit of the formation is c. 8 m thick and shows large-scale cross-bedding with sets up to 3 m thick, and the unit has an r-onlap relationship to the unit below. the upper unit is c. 10 m thick and the beds show mainly planar bedding with some trough cross-bedding towards the top. the lower and middle units may be interpreted to represent deposition during two phases of ramp– flat propagation. gr09 thrust sheet the lønstrup klint formation of the gr09 thrust sheet forms a uniform c. 5 m thick unit with a vertical orientation. the rubjerg knude formation is about 25 m thick and can be divided into a lower and an upper part. the lower part displays variable large-scale crossbedding and the unit has a f-bedding relationship, whereas the upper part forms one large r-onlap succession. the lower part may be interpreted as having been deposited while the gr08 hanging-wall ramp propagated over a footwall flat, whereas the upper part was deposited when the gr09 thrust sheet was displaced up along a 45° dipping ramp during a relatively late phase of deformation. a number of minor horizontal extensional faults are interpreted as foreland-dipping normal faults related to a hanging-wall anticline formed over a hanging-wall ramp during an early or intermediate phase of thrusting. gr10 thrust sheet the lønstrup klint formation of gr10 is very thin, only about 3–4 m thick, in the exposed part of the cliff section. the rubjerg knude formation is about 25 m thick, and the lower part shows a poorly exposed f-bedding relationship. the upper part displays marked r-onlap with a 45° dipping angular discordance to the l/r-unconformity. 117 gr11 thrust sheet the gr11 thrust sheet is irregularly orientated, but is mainly vertical in the upper frontal part. in gr11, the lønstrup klint formation has a uniform thickness of 10 m, consisting of medium-bedded light coloured sand interbedded with thin mud layers situated above the steeply dipping hanging-wall flat. the rubjerg knude formation is generally not well exposed due to sand scree, but has a thickness of about 15 m. gr12 thrust sheet a large part of the gr12 thrust sheet is covered by sand scree, and detailed data from this part of the grønne rende section are limited. the lønstrup klint formation forms a c. 5 m thick unit of sand dominated by thick-bedded turbidites, as is typical of the upper part of the formation. the rubjerg knude formation is more than 20 m thick and displays the typical depositional features of the formation. however, it should be noted that the l/r-unconformity in gr12, as is the case in gr11, has been lifted up to an elevation of 15–20 m a.s.l.; this indicates that these sheets propagated over an upper flat, probably composed of two stacked duplex segments of gr06 and gr08 in the subsurface. gr13 thrust sheet the gr13 thrust sheet is also mainly covered by sand scree at the base of cliff. however, exposures in the upper part of the cliff reveal a rather complex structure. the lønstrup klint formation of the thrust sheet is very thin, and in places only the unconformity is observed; it can thus be difficult to recognise where the stratigraphic unconformity is preserved and where it has been completely replaced by the thrust fault, which is now vertically orientated. furthermore, the rubjerg knude formation has been subjected to superimposed folding. the thickness of the piggyback basin deposits in gr13 is more than 30 m and four units of the rubjerg knude formation are differentiated. unit 1 is about 5–6 m thick, and appears in the frontal and upper part of the piggyback basin; it is characterised by large-scale cross-bedding as well as planar parallel stratification. towards the trailing end of the thrust sheet, the r-onlap in unit 1 grades up into unit 2, which is folded into a set of overturned folds, originally with a horizontal axial plane, but now re-orientated into an upright position; the overturned folds deform the bedding in unit 1. unit 3 is characterised by steeply dipping large-scale foresets that were deposited over the recumbent folds. finally unit 4, about 7 m thick, is mainly planar-bedded and was pushed in front of a hanging-wall ramp of the frontal thrust sheet in the rubjerg fyr section. interpretation of structural development two types of interpretations are considered prior to the further description of the structural development: (1) in order to estimate the displacement, a geometric construction has been made for each thrust-sheet tip subsequently eroded away at the glaciotectonic unconformity (fig. 11), and (2) the position of the l/runconformity below the screes has been constructed from successive approximations (plate 1). interpretation of the extent of the thrust tips was carried out as a triangular construction with a best-fit of the intersection of the unconformity and the thrust plane (see fig. 11). where these surfaces (lines in the 2-d constructions in fig. 11) were obscured, the construction was guided by the assumption that the acute angle is close to 18°, which from experience is a general initial thrust ramp angle. according to this geometrical reconstruction, the average displacement of each thrust sheet is about 70 m. the general impression is that the thrusting of the twelve imbricate sheets in the grønne rende fan-structure was broadly contemporaneous. if the experience from the ulstrup section near the foreland is taken into consideration, one would expect a long coherent thrust sheet initially displaced along the upper flat level at 10 m stratigraphic depth (from the l/r-unconformity). from this detachment level, the imbricate thrust-fault fan propagated with fairly equal spacing, and the thrusting progressed during sedimentation of the rubjerg knude formation. however, there are two positions where the piggyback basin has a thinner depositional fill, namely the basin on gr01 and on gr06/gr07. thus, the first longer displacement along an upper footwall flat took place on the back of gr06 and gr01, subsequently leaving a trailing-end duplex segment below the 10 m detachment level at the rear of gr06 as well as at the end of gr01. when about half of the displacement along the gr06 had taken place, the level of detachment started to root down to the 20 m flat level, such that the trailing duplex seg118 ment of gr06 was free to move along intermediate ramps. thus, if about 20 m of the displacement on each of the gr07–gr13 thrusts is accumulated, it amounts to a total displacement of about 140 m for the gr06 trailing segment (gr06u). it is therefore argued that the gr06u segment started to ramp in the middle phase of thrust propagation, which is reflected in a foreland-dipping tilt in the gr13 basin that created the recumbent folding. the sequence in ramping follows three angular modes: initial ramping along an angle of about 18°, intermediate ramping along an angle of 30°, and final ramping at close to 45°. however, ramping will never be initiated at an angle of 90°, so the problem is how to explain the vertically orientated thrust faults. the first c. 20° ramp is given by the upper ramping, and progressive ramping will result in a 36–45° tilt. this would result in progressive steepening of the tilting from south to north, however, which is clearly not the case. it is suggested, therefore, that the gr01 thrust jumped down to the décollement level at the 30 m stratigraphic depth, and that this caused the reorientation during the final displacement of gr01. thus, when the imbricate fan with ramp angles up to 40° was carried along with the lower gr01 thrustsheet segment (gr01u) towards the 45° inclined frontal ramp, all the thrust sheets were subsequently tilted due to a common megascopic shear. the combination of this large-scale shear tilt and the accumulated ramp steepening is very well illustrated in the gr05 thrust sheet, which was separated from gr04 along a satellite thrust fault. the hanging-wall ramp of gr04 was fixed with a c. 75° steep dip. consequently the hanging-wall ramp of gr05, which was carried piggyback on gr04, ended up in an overturned position (dip of about 75°s). in balancing the grønne rende section, one problem remains to be solved, namely the fate of the trailing-end segment of gr06. however, this is a minor problem compared to the space problem created by the gr01 thrusting along the 30 m décollement level. the trailing-end thrust-sheet segment of gr01u (a duplex sheet created between the 30 m and the 20 m flat level) is considered below under the structural and dynamic analysis of the rubjerg knude fyr and stortorn sections. rubjerg knude fyr section this section is situated below the rubjerg knude fyr, and makes up the highest part of the cliff (see cover illustration). the cliff section below the lighthouse comprises the thickest thrust sheets, which are ramped up to the highest footwall flat level. this means that thrust sheets with a décollement level at a depth of about 35 m are ramped up and thrust along a flat at the 10 m deep level. normal fault structures similar to the brede rende normal fault also appear in the rubjerg knude fyr section. the tip of the prominent thrust sheet in the central part of the section was dropped down into the piggyback basin in front of the thrust fault, and subsequently buried by sediments of the rubjerg knude formation. structural elements related to the hanging-wall ramp are well illustrated in this section. tectonic architecture the rubjerg knude fyr section comprises six relatively thick thrust sheets annotated rf01–rf06. the leading-edge thrust and hanging-wall ramp-and-flat of rf01 terminate all the imbricates in the grønne rende section. to the north, the section is bounded by the footwall ramp of rf06 that coincides with the hanging-wall ramp of the frontal thrust sheet in the storntorn section. the thickness of the thrust sheets is bounded by the 20 m flat level in the southern part, and increases in the northern part down to the 30 m flat. all the thrust sheets in the section are thrust over the lower duplex segment of gr06 and gr01. translation along an intermediate flat is indicated by the dominant position of the l/r-unconformity at 20 m a.s.l. the rf01 and rf02 thrust sheets make up an easily recognisable thrust sheet pair that can also be located on the geological cross-section constructed by jessen (1918; fig. 98). the similarity of the remaining part of the rubjerg knude fyr section to jessen’s cross-section is not so obvious, probably due to the intervening 80 years of cliff erosion and the present poor exposure of the section due to extensive sand scree. the dips of the footwall ramps are about 60°. the angle between the l/r unconformity and the thrusts is about 30°, which means that the thrusts have been rotated about 30° during ramp propagation in the subsurface. this is illustrated by the angular relationship between the l/r unconformity in rf01 and the footwall ramp of rf02. the thrust-fault displacement of rf01 and rf02 ranges from 50–65 m. 119 at the tip of rf01, a 25 m long upper hanging-wall flat is preserved. below this, a 5–7 m thick sheet comprising the top of the gr13 piggyback basin occurs. this upper footwall flat segment of gr13 was displaced an unknown distance (25–50 m) forwards in front of the upper hanging-wall ramp of rf01. the rubjerg knude formation in rf02 is poorly exposed due to sand scree at the base of the cliff, but it is interpreted to be 20 m thick. from the l/r-unconformity up to the overlying thrust fault, the thickness of the piggyback basin is about 40 m. however, it is inferred that a thrust fault situated in the middle part of the basin is responsible for repetition of the rubjerg knude formation, and that a normal fault similar to the brnf, is located in the upper levels of the rf02 sheet. the normal fault is asymptotic, fading out towards the l/r-unconformity in the rf02 thrust sheet. the rf03 thrust sheet is a small sheet with a truncation structure. the nose of this thrust sheet was obviously exposed to normal faulting at an early stage of development (fig. 99). after fault displacement, the rf03 tip was eroded away and the top of the thrust sheet erosionally truncated to form an unconformity, which cut off the sheet at a very steep angle (> 70°). the normal fault at the tip of rf03 is the first of two normal faults displaced down on to the piggyback basin of rf02. in the second phase of normal faulting, the c. 45 m long tip of the rf04 thrust sheet was displaced c. 35 m down a normal fault plane (fig. 100), which had an angle of 70–90° relative to the thrust fault and the l/r-unconformity of rf04. the normal faulting may have been initiated by differential translation of hanging-wall ramps along intermediate flats related to one or more subsurface duplex segments. in a late phase of thrust faulting, the trailing part of rf04 was thrust up over the piggyback basin of rf04/rf03/rf02, which brought the nearly 45° dipping lower hanging-wall ramp of rf04 into contact with the footwall ramp of the displaced tip of rf04. the final phase of fault-bend folding tilted this ramp into a 70°s dipping position. the rf05 and rf06 thrust sheets were thrust up fig. 98. the thrust sheet pair exposed below the rubjerg knude fyr. to the right, the rf01 thrust sheet is thrust faulted along its hanging-wall ramp (rf01hwr) along the footwall ramp in the trailing end of the grønne rende section. rf01 forms the footwall block for the propagation of the hanging-wall ramp of rf02 (rf02hwr). note the elevated position of the l/r-unconformity indicating that the thrust sheets were thrusted up on duplex segments in the subsurface. photograph: july 1999. 120 121 from the 30 m décollement level. the hanging-wall flat of rf06 rests on the upper flat of rf05, where no deposits of the rubjerg knude formation have been recognised. thus the two thrust sheets occur as a block of lønstrup klint formation 60 m thick, separated in the middle by a thrust fault. the l/r-unconformity in the rf06 sheet is located at about 10–15 m above sea level. thus, the rf06 sheet was faulted up on an intermediate flat above the trailing edge of rf05, probably while both were transported along the lower footwall flat on top of the lower trailing duplex segment of the grønne rende section. sedimentary units in the rubjerg knude fyr section, the lower stratigraphic levels of the lønstrup klint formation are exposed, although they are commonly deformed either by mudmobilisation or thrust-fault shearing. the rubjerg knude formation above is poorly exposed, partly due to sand scree derived in part from the formation itself, and partly from the up to 50 m high sand dunes above the cliff. no further description of the formations is given here. facing page – upper: fig. 99. normal faults displacing the top of the rf03 thrust sheet in the rubjerg knude fyr section. note the footwall syncline folded below the hanging-wall ramp at the top of the cliff section. photograph: june 1984. facing page – lower: fig. 100. normal fault (nf) displacing the tip of the rf04 thrust sheet, which prior to normal faulting was thrusted along the hanging-wall flat of rf04 (rf04hwf). photograph: july 1994. fig. 101. thrust-fault brecciation related to the lower hanging-wall ramp and flat in the rubjerg knude fyr section. the brecciation fabric that is typical of the fine-grained sand turbidites and laminated clayey muds (compare with primary sedimentary features in figs 22, 23) was produced by low-angle anastomosing shearing. photograph: september 1985. 122 structures two types of deformation alter the primary sedimentary architecture of the rubjerg knude fyr section. the first type is anastomosing thrust-fault brecciation related to the zone above the hanging-wall ramp with fine-grained turbidites. the second type is the mudmobilisation and mesoscopic-scale polydiapirism, which is common in the clay-rich lower part of the thrust sheets. these deformation types have to be considered in the evaluation of the balance calculation. only the anastomosing thrust-fault brecciation will be further described in this section; descriptions of diapirism are given under the kramrende, brede rende, sandrende and moserende sections. anastomosing thrust-fault brecciation the most significant mesoscopic-scale structure recognised in the rubjerg knude fyr section is related to thrust-fault brecciation in the lower part of the lønstrup klint formation. the brecciation fabric that is typical of the fine-grained turbidites and laminated clayey muds was created by low-angle anastomosing shearing. the shear surfaces and thrust-fault displacements are located in the clay-rich laminae, whereas the segments bounded by the anastomosing fractures consist of silty mud lithologies (fig. 101). the brecciation extends from the lower hanging-wall ramp-andflat up to 10 m above the sole of the sheets. the anastomosing thrust faulting indicates that significant dif ferential movements developed in the subsurface during thrust-fault propagation, which illustrates the nature of the displacement-related shearing and which may account for some of the volume problems arising from construction of the balanced crosssection. interpretation of structural development the compression (shortening) in the rubjerg knude fyr section is about 48%, calculated from the measured length of the section of 270 m (= l 1 ) and the balanced length of about 520 m (= l 0 ). this implies that the thrust sheets in the section have been displaced upwards at the ramp originally situated at the trailing end of rf04 and transported to the position of the footwall ramp of gr13. the lower flat is still the footwall flat on top of the lower duplex segment of the grønne rende section. thus, a considerable amount of lateral translation is apparent in the rubjerg knude fyr section. the appearance of the normal fault with down-faulted noses of the rf03 and rf04 sheets indicates that differential movements, including duplex development of the gr01–gr06 lower segments, may have occurred to create the foreland-dipping features in the rf02– rf03 piggyback basin. the interesting structure in rf05 is the lower hanging-wall ramp, which has to correspond to a footwall ramp at the trailing end of rf04. the footwall ramping probably also included propagation along a footwall ramp of a lower duplex segment of rf04 (rf04u). the present orientation of the rf05 lower hangingwall ramp is more or less vertical, indicating three steps of ramping. the final tilting was due to the ramping of the upper footwall ramp of gr13, the middle phase of ramping was up along the footwall ramp of the piggyback basin in the normal fault displaced rf04, and the initial ramping was probably a complex propagation over several smaller ramp-steps that separated rf06 from rf05. the uppermost part of rf06 shows a marked topography, indicating that at an early phase of deformation it was elevated up to a level of erosion, before subsequent sedimentation. this sedimentation was probably of relatively short duration before over-thrusting of the stortorn section trapped the piggyback basin. stortorn section stortorn is the name of the very steep and muddy cliff in the central part of the rubjerg knude cliff section. on old drawings of the beach and the coastal cliff, stortorn is depicted as a steep, wild looking castlelike cliff in the distant horizon, emphasising the romantic scenery of this remote place (e.g. engraving by c. neumann 1884, reproduced in vendsyssel nu og da, 1981). in recent times, the cliff has also been the location of major landslides, which in some cases travelled more than 100 m out into the sea. in general, the sea reaches close up to the vertical cliff, and due to the muddy and slippery cliff surfaces and the clay pavement in the zone of breakers, it is the most difficult place to pass along the coast. in the stortorn section, the deepest level of thrusting occurs where the décollement zone is located at a depth of 40 m stratigraphically below the reference surface of the l/r-unconformity, which is about 45 m 123 below sea level. from this deep level, the thrust sheets were elevated up to the exposed position in the cliff section. coinciding with this, the thrust sheets contain the deepest levels of the stratigraphy, and beds of marine and glaciomarine clay can be identified by the occurrences of arctic marine fossils. moreover, the stortorn section contains the key features for understanding the structural and dynamic problems of the adjacent rubjerg knude fyr and grønne rende sections. the key features are flat-lying duplex complexes formed by ramping up of the relatively long, lower thrust-sheet segments onto a high flat level. this is reflected in an elevation of the l/r-unconformity up to a height of c. 40 m above sea level in the cliff section. tectonic architecture the stortorn section is divided into ten thrust sheets annotated st01–st10. the southern boundary of the section is the frontal hanging-wall flat of st01, which coincides with the footwall ramp of the rf06 thrust sheet in the rubjerg knude fyr section. the northern boundary is the trailing footwall thrust of st10 along which the frontal hanging-wall ramp of the moserende section was thrust. the southernmost three thrust sheets form a separate group of high-level thrust sheets. the l/r unconformity is here situated at an elevation of 35–40 m above sea level. the central part of the section is formed by a series of thick thrust sheets of clayey, partly mobilised, mud, which are situated above hidden duplexes in the subsurface. this complex extends about 150 m along the cliff section at stortorn. in the northern part of the section, upright mud diapir-dominated thrust sheets occur with complexly developed piggyback basins. st01 thrust sheet the southernmost thrust sheet in the stortorn section (st01) is wedge-shaped, c. 160 m long, with an initial 25° dip of the frontal hanging-wall ramp. the footwall ramp (fr06 trailing edge) dips at about 45°, which creates a problem in the balancing. it is obvious that a splint (or horse) corresponding to a triangle with an acute angle of 20° must be hidden somewhere in the deeper structure. the trailing end of st01 is 30 m thick. however, the lønstrup klint formation is deeply eroded in the central part of st01, which truncates the l/r-unconformity. st02 thrust sheet the st02 thrust sheet was also elevated to a height of 35–40 m in the cliff section. the hanging-wall ramp is vertical and forms a right-angle with the horizontal bedding in the piggyback basin of st01. this indicates that final up-thrusting of the frontal hanging-wall ramp took place in an upright position during sedimentation on the back of st01. the st02 thrust fault was rotated at least three times before propagation up along the footwall ramp of st01. the initial dip of this ramp was relatively steep, about 40°, as indicated by the angle between the thrust fault and the bedding in the lønstrup klint formation of st01. furthermore, the problem related to the change in ramp angle recurs. in the balancing of the thrust structure, a splint volume must be calculated for, corresponding to the triangle created by the initial thrust angle and the final steeply inclined thrust fault. st03 thrust sheet the st03 thrust sheet is not elevated as much as st01 and st02; the l/r-unconformity is only situated at about 30 m above sea level. the lønstrup klint formation was deeply eroded before sedimentation in the piggyback basin was initiated, probably due to marked relief during ramp propagation. st04 thrust sheet the lowest position of the l/r-unconformity in the st04 thrust sheet is about 13–15 m above sea level, which demonstrates a shallower level of ramping than in st01– st03. the 45° steeply dipping thrust fault between st03 and st04 roots down to the décollement zone and propagated up along the footwall ramps of the subsurface duplex segments beyond st01–st03. the dif ference between initial and final angle of ramping also created a balancing problem for st04. the frontal part of the hanging-wall ramp of st04 only had a dip of 18–20°, whereas the footwall ramp now dips at c. 45°. therefore a splint (with an area of 630 m2 in the cross-section) is envisaged in the subsurface. the st04 ramping over this splint is interpreted as the rea124 son for the elevation of the l/r-unconformity up to c. 15 m above sea level. st05 thrust sheet the st05 thrust sheet is one of the most important structures, not only in the stortorn section but also in the rubjerg knude glaciotectonic complex as a whole. it involved thrusting of the deepest décollement level, which introduced a complex framework due to the large number of duplex segments involved. the st05 thrust sheet was thrust up along the footwall ramp of st04. the dip of the thrust fault increases from 35° at the beach level to 70° at the top of the cliff section. the tip of the thrust sheet was finally displaced horizontally over the top of the piggyback basin of st04. the thrust displacement along the hanging-wall thrust fault is estimated to be in the order of 90 m. the lønstrup klint formation of st05 is dominated by mud diapirism and polydiapiric structures up to 5 m in vertical scale. in the middle part of the thrust sheet, where the thickness of the initially wedgeshaped frontal part was 20 m, a deeply eroded trough was formed. the l/r-unconformity on the northern flank of this erosional depression is situated nearly 50 m above sea level. this is about the highest elevation of the reference surface, and was caused by thrust duplication of the thrust sheet in the subsurface duplex complex. st06 thrust sheet the st06 thrust sheet is poorly exposed and mudmobilisation and internal diapirism obscure primary structures. the thickness of the thrust sheet is up to 40 m, measured from the frontal hanging-wall thrust up to a small pocket of rubjerg knude formation sand that forms the remnant of a piggyback basin. the frontal thrust is drawn with some uncertainty, because a large part of it is penetrated by diapirism intruding from the back of st05. a minimum displacement of 40 m is inferred, which is incorporated in the modelling of the balanced cross-section. this implies a rather complex structural assemblage of the subsurface lower duplex segments of st05, st06 and st07. st07 thrust sheet the st07 thrust sheet is nearly vertically orientated with a frontal ramp rising from 70° to vertical, along which a displacement of 28 m is estimated to have occurred. the l/r-unconformity surface is also steeply dipping, and is even overturned at the top. the lower part of the thrust sheet is mainly covered by scree, but it is possible to trace the line of the unconformity down to the level of the beach in the crosssection. this implies that the thrust sheet has not been thrust up to be displaced along an intermediate ramp but is only tilted due to the main ramping, first along the thrust fault of st06 and finally on its own hanging-wall ramp. the ramping is also reflected in the deposition in the piggyback basin where three superposed angular discordances are recognised. the deposits in the piggyback basin are c. 15 m thick; at the base of the succession, the r-onlap starts with an angle of 45° and terminates with a 90° angle, indicating deposition in the basin while the thrust sheet was vertically orientated. in the uppermost bed, minor slump folds are present, indicating the effect of the st08 thrust nose approaching from the north. st08 thrust sheet a double ramp synclinal structure, similar to the one occurring in the piggyback basin of st05, is recognised in thrust sheet st08. the l/r-unconformity incises through the lønstrup klint formation and down into the thrust-fault surface of the hanging-wall ramp. the lateral distance between the ramps is only about 25 m and the basin is less than 10 m deep. the sediments in this piggyback basin show large-scale trough cross-bedding accentuated by synclinal folding. the rubjerg knude formation covers a feature that represents the erosional remnants of a detachment anticline on the northern limb of the basin. on the north side of this structure, the initial stratification above the l/r-unconformity shows clear r-onlap, corresponding to the inclination parallel to the tilt of the lower ramp. the unconformity is elevated up to 20 m above sea level, indicating that the st08 thrust sheet was lifted up by at least two subsurface duplex segments. these hidden segments are annotated st08u1 and st08u2. the displacement along the footwall ramp of st08 is estimated at 78 m, mainly along the thrust fault dipping at 30°. 125 st09 thrust sheet the st09 thrust sheet is a c. 30 m thick sheet bounded by a 60° dipping hanging-wall flat thrust up onto the footwall ramp of st08 back and the 60° dipping footwall thrust fault of st10, which truncates the irregular structures in the upper part of the thrust sheet. the displacement along the hanging-wall ramp is estimated at c. 60 m. although the lønstrup klint formation in st09 is characterised by internal diapirism, the features of a hanging-wall anticline can be recognised at the top of the cliff section. during translation of two or more footwall ramps, an irregular synform formed and created the depocentre of a piggyback basin. the l/r-unconformity is here elevated to 5–10 m above sea level, corresponding to propagation up onto the st08u2 duplex segment. st10 thrust sheet the st10 thrust sheet is about the same size as the st09 sheet, and also has steeply dipping bounding thrust faults. the most remarkable structure in the st10 thrust sheet is the structural complexity of the piggyback basin. the l/r-unconformity forms an isoclinal recumbent syncline, with the upper limb formed by the mud-mobilised lønstrup klint formation, and above this a minor synclinal trough appears. this structure is best described as a detachment anticline, which developed into a diapir with a reverse fault displacing the northern limb of the structure into a mushroomshaped structure, similar to the diapir in the sandrende section. at a late stage of thrusting, the piggyback basin was rotated 60° and the diapir-developed detachment anticline collapsed into the recumbent strucfig. 102. isoclinal upright anticline formed in the lower part of the lønstrup klint formation in the frontal part of the stortorn section. the right limb of the anticline constitutes an imbricate duplex formed by connecting thrust-fault splays (white dot-and-dash lines). the fold is interpreted as a hanging-wall anticline developed during fault propagation and successive imbricate stacking (compare with fig. 59). photograph: august 2001. 126 ture. the l/r-unconformity is elevated up to 15–18 m above sea level indicating a ramping of the st08u duplex segment as well as the trailing segment of st09. an estimated displacement of 73 m along the hanging-wall ramp of st10 still leaves some subsurface segments to be balanced in the structure below the moserende section to the north. sedimentary units the most important sedimentary feature in the stortorn section is the exposure of the stortorn formation, which is the lowermost stratigraphic level involved in the rubjerg knude glaciotectonic complex. the stortorn formation is located at the lower hanging-wall ramp of the st05 thrust sheet, where it forms a duplex segment about 3–5 m thick at the base of the cliff. to date, it has not been possible to measure a sedimentological log of the formation at this locality, partly because the formation is strongly sheared by anastomosing fractures, and partly because landslide activity precludes more detailed stratigraphic description. structures the contact between mobilised, intrusive mud and stratified mud has been observed in many places, but this characteristic is better illustrated in the moserende section (see below). anastomosing thrust faults and tectonic breccias occur commonly in the st05 and st06 thrust sheets. however, due to the difficult field conditions, detailed investigations have not been carried out. in the dark clayey mud of thrust sheet st01, an upright nearly isoclinal anticline has been observed (fig. 102). this fold is considered to represent a hanging-wall anticline that was subjected to an advanced stage of deformation during ramp propagation. this stage compares well with the model of duplex formation described by mitra & sussman (1997), in which the growth of imbricates derived from successive connecting splays results in steepening of antiformal stacks formed by fault-propagation folding of duplexes. successive growth of duplex elements corresponds well with the interpretation presented below. interpretation of structural development the cross-section (plate 2b) provides a model for the structures below the frontal part of the stortorn section, which requires four duplex segments forming a duplex complex, on top of which thrust sheets st01, st02 and st03 have been thrust along the footwall flat. the structural interpretation of a duplex stacking of subsurface segments below the frontal part of the stortorn section is based on two lines of evidence: (1) the missing balance of the lower segments related to the grønne rende section, and (2) the high elevated position of the l/r-unconformity in this part of the section. thus the lower duplex complex in the frontal part of the section is interpreted to represent stacking of the trailing lower segment of gr01, although an alternative differential duplex-segment displacement is also possible. from geometric considerations, it is evident that the 20 m and 30 m décollement levels must have been pervasive throughout the proximal part of the thrust structure. thus the duplex in the stortorn section consists of segments c. 10 m thick. mobilisation and the internal polydiapirism have obscured the boundaries of these segments, which are the lower and intermediate footwall and hanging-wall flats respectively. however, in a model for reconstruction, these volumes are regarded as solid thrust sheet, i.e. the duplex segments (represented by annotated areas in plate 2). in the description and solution of the structural problem related to differential thrust faulting of the lower duplex segments, three main types of fault-bend-folded segments are distinguished (fig. 103). fig. 103. schematic illustration of the three types of fault-bend folding of duplex segments. type 1 is referred to as an l-structure, type 2 as an s-structure and type 3 as a g-structure (g chosen due to similarity with the greek capital letter gamma (γ)). the footwall ramp dips at about 45°, the shortening between the underlying footwall ramp and the overlying hanging-wall ramp is 43%, the initial length (l 0 ) of the thrust sheet is c. 100 m, compared to a thrust-sheet thickness of 40 m, and a thickness of 10 m for the individual duplex segments. 127 1. a duplex segment with one part of the segment resting on the lower flat and the other parallel with a ramp (l-structure). 2. a duplex segment which has the trailing part parallel with the lower flat, the intermediate part parallel with a ramp, and the frontal part parallel with the upper flat, thus giving the shape of a s. 3. a duplex segment with the trailing part of the segment located parallel with the ramp and the frontal part parallel with the upper flat (γ-structure). the stacking of the duplex below the st01, st02 and st03 thrust sheets started when the lower segments were thrust up into the first type duplex bend during ramping towards the footwall ramp of rf06 (trailing end of the rubjerg knude fyr section). this probably marked the end of the lateral translation along the lower flat levels and the initiation of stacking along steeply dipping thrust faults. this resulted in the rotation of all the previously formed structures and created the odd trough structures in the double ramp synclinal troughs. the final up-thrusting along steeply dipping thrust faults, which occurred contemporaneously with the uppermost sedimentation in the piggyback basins, was probably also contemporaneous with the initiation of ramping of st05 from the lowest level. one way to demonstrate this is to focus on the thrustfault development of st04. the st04 thrust fault acted as the ramp that pushed on the trailing end of the duplex below st01, st02 and st03. when the push on this footwall ramp ended and the st04 thrust sheet was displaced up over the footwall ramp of st03 it resulted in a shortening of 60%. the balanced length of st04 is c. 200 m. thus, the deep level st05 ramp must be responsible for removing the 120 m lower long segment originally situated below st04. therefore about half of the st04 thrust sheet also involves thrusting down to the 40 m décollement level. in the model, this st04u segment was up-thrust to form a first type of duplex-segment structure as the first lower-segment imbricate in the subsurface of st05. however, it should be appreciated that a whole unit of the thrust segments between the 20 m flat level and 30 m décollement level has to be incorporated in a differential thrust model. the simplest model for this is to dissect the lower duplex segments into sheets with an average length of c. 100 m. with each segment bend in a type 2 ramping and with an equal distribution of the frontal and trailing part on the upper and lower flat, a series of double ramp synclines would be created – comparable to the piggyback basins seen in st10, st09, st08 and st05. similar basins may have existed in st07 and st06, but, if present, were removed by glaciotectonic truncation. one of the central problems in describing the dynamic development of the rubjerg knude glaciotectonic complex is understanding the formation of the lower ramp below the st05 thrust sheet. it is known that the hanging-wall ramp is displaced up along the st05 thrust fault to be exposed in the cliff section at stortorn. however, a central question is – where was the footwall ramp for the lower hanging-wall ramp of st05 situated? according to the balanced cross-section, the st05 ramp should be situated about 7800 m from the frontal ramp in the ulstrup section and the footwall ramp for the lower st05 ramping should be situated on the far side (north) of the lønstrup village. however, this is not the position of the st05 ramp. the distance to a hidden footwall ramp can only be fixed relative to the displacement in front of the lower ramp when it was activated during the change of décollement level from the 30 m level down to the 40 m flat level. the relative displacement on the st05 thrust fault is c. 90 m, measured from the tip of the thrust sheet along the hanging-wall ramp-and-flat down to the lower décollement surface. the main problem is related to the compression documented south of the stortorn section. when this is considered in the balanced crosssection, it gives the geometric point for the hangingwall ramp 7800 m from the frontal ramp in the ulstrup section. however, the distance in the rubjerg knude cross-section from this ramp to the central part of the stortorn section is only 3800 m. the solution to this problem is that the ramp was first formed after all the former translation in the higher flat levels had passed. to understand this, one has to imagine that the upper part of the lønstrup klint formation at grønne rende (c. 3800 m from the frontal ramp in the ulstrup section) was originally situated above the stortorn formation at the st05 lower ramp. however, the ramp was first formed when the rubjerg knude fyr section was displaced towards the grønne rende section which itself was compressed against the stenstue and sandrende sections, which were all displaced over the initial position of the brede rende and kramrende sections. only then was the st05 lower ramp activated, and the remaining northern part of the rubjerg knude glaciotectonic complex was displaced along the lowermost décollement zone. 128 moserende section between rubjerg knude fyr and mårup kirke, some small peat-bogs occur in depressions in the dune landscape. this area of bog-filled depressions formerly extended to the west, and the present cliff section intersects one of these bogs where peat (martørv) is exposed in the uppermost part of the cliff, similar to the situation in the martørv bakker section. a former gully here was named moserende, and the name is adopted here for the section north of the stortorn section. the most impressive feature in the moserende section is the syntectonic evolution of the piggyback basins during polyphase thrust propagation (fig. 104). unusual sedimentological features are developed in the rubjerg knude formation, reflecting the tectonic deformation, notably structures described as fissure strata (sjørring 1977). these are thin sedimentary beds occurring as discordantly incised wedges in groundfrozen sediment, here present in the growth-fault synclines related to the piggyback basins (fig. 105). tectonic architecture the moserende section comprises 13 relatively thick thrust sheets annotated mr01–mr13. in the section, three larger piggyback basins are preserved, one in the frontal, southern part and two in the northern part of the section. the l/r-unconformity surface at the base of the piggyback basins was elevated to various levels in the cliff section, reflecting the differentiated type of ramping throughout the section. the leading-edge thrust in the moserende section is the hanging-wall ramp of mr01, which coincides with the footwall thrust fault on the back of the st10 thrust sheet. to the north, the section is bounded by the hanging-wall thrust at the base of the c. 40 m thick mk01 thrust sheet, which forms the southern front of the mårup kirke section. the mk01 thrust sheet was thrust up along a steeply dipping footwall ramp and subsequently displaced over the upper footwall flat on top of the piggyback basin of mr13 in the northernmost part of the moserende section. it should be noted that the main frontal part (topfig. 104. the piggyback basin of the mr02 thrust sheet. a sequentially developed growth-fault footwall syncline was formed below the hanging-wall ramp of mr03 (mr3hwr ). note the fissure strata cross-cutting the bedding (arrow) in the growth-fault footwall syncline. photograph: june 1984. 129 most part) of the thrust sheets shows a marked drag and truncation due to the formation of a glacitectonite on top of the cliff section. mr01 thrust sheet in the cliff section, the mr01 thrust sheet forms a massive unit of mobilised, structureless grey mud. the exposed thickness close to the beach level is nearly 30 m. the hanging-wall thrust dips at about 60°n, and the unconformity surface, which dips at about 45°, is elevated c. 15 m above sea level. this indicates that the mr01 thrust sheet is situated above two lower duplex segments. according to the balancing, these segments constitute a lower segment of the mr01 thrust sheet (mr01u) and a segment originating from the lowest, northern part of the stortorn section. the mr01u segment, which exists as a consequence of the estimated c. 46 m displacement along the 60° tilted frontal hanging-wall ramp, is separated into two differently displaced segments in the balanced cross-section. this is a feasible explanation but not the only one of several possible solutions for the displacement structure in the subsurface, which include differential lateral displacement along each lower 10 m level as well as mud diapirism. fig. 105. detail of the fissure strata indicated in fig. 104, illustrating that climbing ripple cross-laminated sands were deposited in the initially horizontal wedge-shaped fissure extending out into the growth-fault syncline deposits. photograph: june 1984; notebook (c. 18 cm long) for scale. 130 mr02 thrust sheet the mr02 thrust sheet is nearly 40 m thick and is dominated by mud mobilisation and internal chaotic structures reflecting polydiapirism and internal flow. the sheet is divided internally by a thrust-fault zone with dif ferential thrust movements, which could be interpreted as a separation of the sheet into two individual thrust sheets. however, as the segments are not separated by a piggyback basin they are regarded as a single amalgamated sheet. the displacement along the hanging-wall ramp is the same order of magnitude as for mr01 (c. 47 m) and the ramp is divided into an upper low-angle part with an initial dip of only 20° and a lower steeply dipping part with an initial ramp-angle of about 45°. due to subsequent rotation, the upper part of the ramp is now orientated vertically while the lower part has a steep listric dip to the south. the footwall ramp (upper part of mr01) has an initial dip of 45°, which creates a space problem in balancing the section and makes it necessary to introduce a splint segment between mr01 and mr02 in the subsurface. the splint was probably sheared and squeezed out and is likely to have been included in the general mud mobilisation. however, it is included in the balanced profile in order to deal with the ramp-angle-space problem (plate 2). the geometry of the l/r-unconformity at the top of the lønstrup klint formation in the mr02 thrust sheet is very irregular with a peculiar c. 8 m high obstacle. this is very similar to the structure in the st08 thrust sheet in the stortorn section. it was probably formed by a detachment anticline on the northern flank of the piggyback basin in the external part of the mr02 thrust sheet, where it was subsequently buried by the rubjerg knude formation sand. the unconformity is elevated about 5 m above sea level, indicating that the mr02 thrust sheet has only stepped up one level of the lower segments from where it is bent up along its hanging-wall ramp. the detachment anticline was probably formed during the ramping of this lower segment. mr03 thrust sheet the rubjerg knude formation in the piggyback basin on top of mr02 and mr03 is here envisaged as a single large basin situated in the frontal part of the moserende section. the basin extends about 70 m along the cliff section. three upright standing peaks represent the tips of three small thrust sheets that disturbed the basin by small displacements. these thrust sheets represent imbricates in the uppermost part of the mr03 thrust sheet. since the main part of the mr03 thrust sheet can be viewed as one large sheet subjected to a single mode of displacement, the three imbricates are referred to as mr03a, mr03b and mr03c. the l/r-unconformity in mr03b and mr03c can be traced down below sea level, indicating that the main part of mr03 was displaced along the lower décollement level prior to the displacement up along the 45° dipping frontal footwall ramp. however, the frontal hanging-wall ramp of mr03 is now vertical. it is only necessary to tilt the initial ramp on another 45° dipping ramp to achieve this, and although it is a very steep inclination for ramping, there are no obvious reasons for introducing more ramps. the steep ramp angle of the mr03b thrust forms part of the same framework. the initial dip of the mr03b hanging-wall ramp was 18°, and the thrust fault is now vertically orientated due to the ramp-bending mentioned above. the displacement relative to mr03a is only about 10 m and the sand beds of the rubjerg knude formation were folded in a footwall syncline of mr03a during the hanging-wall thrusting of mr03b. mr04 and mr05 thrust sheets the mr04 and mr05 thrust sheets are closely related and only separated from each other by a relative displacement of about 25 m along the hanging-wall thrust of mr05. in contrast, the displacement along the mr04 hanging-wall thrust is about 80 m. both thrust sheets are dominated by mud diapirism, structures that may have originated as one large diapir that was only displaced by the late mr05 hanging-wall ramp. the thickness of the lønstrup klint formation in the thrust sheets is up to 30 m in the cliff section, and the height of the diapir is 15 m. the diapir has characteristic intrusive contacts with the upper and frontal part of mr04 (fig. 106). it is evident that the final thrust displacement post-dates the diapirism, and the very steep thrust angle indicates that the ramping is rooted in the deepest levels of the section. the l/r-unconformity is elevated up to 10–12 m above sea level. from this it is inferred that the thrust sheets were lifted up on the lower segments of the mr03 and mr04 thrust sheets, although most of the lift is related to the ramping on the steep thrust faults. the thickness of the 131 rubjerg knude formation in the piggyback basin of the thrust sheets is about 15 m. mr06–mr08 thrust sheets the main feature of the mr06 and mr07 thrust sheets is that they are lifted relatively high up in the cliff section, such that only a small part of the piggyback basins are preserved. the ramping of mr06 and mr07 is about 22 and 13 m, respectively, corresponding to one level of elevation of the foremost mr06 thrust sheet. the ramping is interpreted to have been a stepwise progression up over the trailing lower part of mr05, which is ramp-bent over the lower segment of mr04. the displacement of mr06, mr07 and mr08 on each hanging-wall thrust fault is about 40 m, indicating that the displacement is of the order of the distance down to the décollement surface. the general impression is that mr06, mr07 and mr08 initially formed one large thrust sheet, which was stepwise separated during dif ferential thrust movements. this dif ferential thrusting moved mr06 to the highest position, whereas mr08 was left in the trailfig. 106. mud mobilisation and diapirism in the lønstrup klint formation in the central part of the mr04–mr05 thrust sheets of the moserende section. photograph: may 1985. fig. 107. mobilised mud (lower left) in the lower part of the lønstrup klint formation intruded into the bedding of the formation. the mobilised mud probably formed a viscous liquid that facilitated the gravity-spreading deformation mechanism. photograph: june 1984; staff divisions are 20 cm. 132 ing part with its hanging-wall flat still resting on the lower décollement surface. this is implied by the elevation of the l/r-unconformity, which can be traced down to a horizontal orientation about 5 m below sea level in mr08. in the frontal part of the mr08 thrust sheet, a very well-developed intrusive contact of a diapir is exposed (fig. 107). it is evident that the process of thrust faulting was facilitated by the buoyancy and lubricating effects of the water-saturated mud. in addition, the mobilised mud had the effect of pushing the thrust sheets from the rear during the gravity spreading process. mr09 thrust sheet the mr09 thrust sheet is a relatively thick thrust sheet. the angle between the main part of the hanging-wall ramp and the bedding in the lønstrup klint formation within the sheet is c. 30°, and the displacement is estimated to be about 58 m. the l/r-unconformity is elevated to about 7 m above sea level, and the mr09 thrust sheet must be considered to have been displaced along the rear part of the lower segments in the section. mr10 thrust sheet the mr10 thrust sheet has the same characteristic shape as the st08 and mr02 thrust sheets, with an obstacle interpreted as a detachment anticline. the final orientation of the hanging-wall ramp is rather steep (c. 70°) implying rotation of an initially steep ramp (c. 35°). this corresponds well with the angle between the bedding in the piggyback basin of mr09 and the mr10 hanging-wall ramp. in the cliff section, the lønstrup klint formation within the thrust sheet is up to 35 m thick. due to uncertainties in reconstruction of the frontal part, the displacement is estimated to be between 30 and 65 m. the l/r-unconformity can be traced down to about 5 m below sea level, indicating that the main part of mr10 rests on the lower décollement surface. the rubjerg knude formation of mr10 is about 25 m thick. it contains a c. 12 m thick lower unit with bedding dominated by r-onlap. above this follow three units, each developed as growth-fault footwall synclines. to obtain the rather large accumulated thickness of deposit in the piggyback basin, as well as folding the three synclines, the order of displacement is more likely to be 65 m than 30 m. this order of displacement also assumes that the frontal part of the thrust sheet extended nearly 50 m further ‘up in the air’ before being removed by erosion. the final implication is that the upper piggyback basin above the anticlinal obstacle was deposited in a syntectonically deeply eroded depression. mr11 thrust sheet mr11 is a small thin thrust sheet with a displacement of 55 m along the hanging-wall ramp. the thrusting of mr11 is another example of a thrust sheet requiring the formation of a splint in the subsurface. this is due to the low angle of the initial frontal ramp (only 15–18°), whereas the ramp angle between the hanging-wall ramp and the original bedding in the piggyback basin of mr10 is 40–45°. the splint was probably trapped as a triangular prism along the 45° dipping ramp, just below the beach surface. the l/r-unconformity is elevated to about 7 m above sea level. the model for balancing this thrust sheet indicates that it had a lower hanging-wall flat situated at the 30 m level (below the l/r-unconformity). from this level, it ramped up to the 20 m level onto a footwall flat on mr10, and finally developed an internal duplex of its own lower segment, which separated into mr11u 1 and mr11u 2 (see plate 2). the piggyback basin of mr11 consists of only a 10 m thick unit of the rubjerg knude formation. the sand beds show a f-bedding relationship to the l/runconformity with a weak tendency for d-onlap. mr12 thrust sheet the exposed part of the lønstrup klint formation in the mr12 thrust sheet is about 15 m thick and was displaced c. 58 m along a hanging-wall ramp dipping 38°. about half of the lønstrup klint formation here constitutes mobilised mud, without showing any marked tendency towards diapirism. the rubjerg knude formation is typically c. 15 m thick, with the exception of increased thicknesses in growth synclines, and the mr12 thrust sheet is regarded as only 30 m thick. this leaves another nearly 60 m long lower segment to be added to the trailing end of the mr10u duplex segment. the initiation of the mr12 thrusting started relatively early compared to the thrust sheets in front (to the south) and behind. this is based on a 133 consideration of the thickness of the rubjerg knude formation in mr11, which only reached a thickness of about 10 m before it became trapped by the mr12 thrust sheet. in the cliff section, the mr12 hangingwall flat (related to the 20 m flat level) is positioned on the footwall flat of the mr11 thrust sheet. the l/runconformity is situated about 5–6 m above sea level, which is compatible with the lift of the thrust sheet up onto the footwall flat of its own lower segment (mr12u). the rubjerg knude formation of mr12 was deposited in one of the large piggyback basins in the proximal part of the moserende section. the width of the basin is about 45 m and the accumulated thickness of the sand succession is c. 23 m. four units can be differentiated in the rubjerg knude formation of mr12. the lowest of these is about 10 m thick, displays fbedding, and corresponds well to the lower part of the rubjerg knude formation in other parts of the section. above this follows a trough-shaped unit folded in a gentle syncline. a new trough-shaped unit, which is folded into a tight footwall syncline, truncates the northern limb of this syncline. finally these two growth synclines are truncated by the upper sub-unit; the latter is mainly f-bedded, except for the northern part which is dragged into a footwall syncline below the hanging-wall thrust of mr13. mr13 thrust sheet the lønstrup klint formation of the mr13 thrust sheet forms an upright, wedge-shaped feature, which was displaced about 42 m up along a relatively steep (70° dip) hanging-wall ramp. this thrust surface has a remarkable curved shape, which is interpreted to be the result of erosion in the ramp caused by water flow contemporaneous with the deposition of the two growth synclines in the external part of mr12. similar features have been observed in places further to the south, but this is one of the best-developed examples. the erosion can be determined to have taken place in the interval between deposition of the lower 10 m thick unit of the rubjerg knude formation subsequent to c. 20 m ramping and before the final displacement along the ramp and deposition of the uppermost unit in the mr12 piggyback basin. the l/r-unconformity is situated 5–6 m above sea level. this is close to the elevation of the l/r-unconformity in mr12, and the thrusting follows a similar development style with ramping and displacement along the flat of the lower segment of the trailing end of the thrust sheet in front (to the south). the section balance requires a lower segment of the mr13 thrust sheet (mr13u) in the subsurface, bounded by the flats at the 20 m and 30 m levels (plate 2). the mr13u segment accumulated on the mr10u segment leads to a high ramping of the thrust sheet to the north, as will be demonstrated in the following section. the large piggyback basin in the moserende section is represented by the c. 20 m thick succession of the rubjerg knude formation in mr13. it is nearly 60 m wide and the main part is planar parallel bedded (f-bedded); only in the uppermost, rear part of the thrust sheet does a single footwall syncline appear. the external part of mr13 is therefore thought to have been deposited during a relatively long period of transport along a flat, contemporaneous with the ramping that had started in the thrust sheets in the distal part of the section. sedimentary units in the moserende section, only two formations involved in the thrust-fault deformation can be studied. the lønstrup klint formation forms the lower part of the thrust sheets, and the rubjerg knude formation forms the fill of the piggyback basins above the l/r-unconformity. lønstrup klint formation the lower part of the lønstrup klint formation typically consists of clayey mobilised mud. about twothirds of the thrust sheets comprise mobilised mud, which commonly developed into diapirs, rising from the sole thrust up into the formation, where they often truncate bedding with intrusive contacts (fig. 107). the upper part of the formation is composed of thinto medium-bedded sandy turbidites interlayered with silty mud (fig. 27). in the majority of layers, the mud and fine-grained sands have been disturbed by balland-pillow breccias or small polydiapiric water-escape structures. during the main thrust faulting, these structures were superimposed by a dense framework of smaller thrust faults (fig. 27). 134 rubjerg knude formation the rubjerg knude formation has a maximum thickness of 25 m in the moserende section; such thicknesses are rarely attained, however, either due to overthrusting that sealed the deposits in the piggyback basins before accumulation of this thickness, or to erosion of the upper part of the formation after deposition and deformation. in general, deposition was initiated with a unit of f-bedded sand c. 10 m thick. above this are two or three units showing r-onlap and locally foreland-dipping large-scale cross-bedding (d-onlap). the uppermost part typically shows r-onlap. structures in the moserende section, three types of structural elements are described and discussed: (1) diapirs, including mud mobilisation, (2) mesoscale thrust faulting, and (3) growth-fault footwall synclines, including fissure strata. the existence of frozen sand clasts and frost wedges that testify to the ground-frozen condition of some of the sediments in the rubjerg knude glaciotectonic complex has already been mentioned. further evidence for ground-frozen conditions is seen in the presence of fissure strata (sjørring 1977). a fissure stratum is a layer of sand deposited horizontally in a fissure that discordantly cuts into a package of sediment (usually meltwater sand) (fig. 105). the implication of the occurrence of fissure strata is that not only was the host sediment affected by (glacio)tectonic deformation prior to fissure incision, but also that the sediment must have been frozen so that the fissure cavity did not collapse during deposition of the fissure strata. as the fissure strata form wedge-shaped sand layers, they have also been referred to as kilelag (wedge-layers in danish; berthelsen 1975). in the moserende section, the fissure strata document an intermediate phase of syntectonic deposition, as they have been tilted into a vertical position. diapir structures in the moserende section, the zone above the hanging-wall ramps and flats often comprises mobilised mud (figs 106, 107). the mud-mobilisation and related diapirs dominate in the thrust sheets from the intermediate hanging-wall ramp and towards the trailing end of the sheets. the diapirs are irregularly developed, mainly related to intrusive migration laterally into the bedding. in the example shown in fig. 107, the intrusive mobilised mud has a contact rim of segregated sandy mud that forms the contact to the truncated bedding. in the thickest thrust sheets, the larger diapirs rose from the hanging-wall thrust fault up to 15 m above the thrust-fault surface (fig. 106). in the construction and calculations of the balanced cross-section (plate 2), the mobilised mud and polydiapirism create a problem of volume preservation relevant to the approximation and evaluation of the reliability of the balanced model. however, the volumes are not lost but only reorganised and may therefore be treated as part of the thrust sheets and duplex segments. the mobilisation is dominantly developed along the lower part of the thrust sheet, from where the mud intruded the upper part of the lønstrup klint formation of the thrust sheets. it is evident that the thrust sheets were carried on the mobilised mud, which with its high water pressure facilitated the displacement of the sheets. pedersen (1987) described a model for this process, and with minor modifications this is still considered to be valid (fig. 4). the model also implies that the muddy liquid formed a pressure agent in the gravity-spreading dynamics, which pushed the thrust sheets forward towards the distal part of the thin-skinned thrust-fault complex. thrust faults in the structural cross-section, only the thrust faults identified as carrying the major displacements are outlined (plate 1). a number of smaller thrust faults and bedding-parallel contractional faults that occur in the rubjerg knude glaciotectonic complex are therefore not included in the cross-section. one example of these less significant faults was described in the rubjerg knude fyr section (fig. 101). in the moserende section, similar thrusts occur, and were recorded during detailed logging of the upper part of the lønstrup klint formation in the mr03 thrust sheet (fig. 27). within a 7 m thick succession, ten minor thrust faults have been recognised. the base of the succession is the hanging-wall ramp of mr03b. along the base of the succession, 0–1.5 m above the hanging-wall ramp, the clayey mud is cataclastically brecciated by anastomosing shear fractures. in the overlying part of the formation, the bedding-parallel thrust faults occur with a spacing of 0.5–1.5 m, concentrated in the muddy 135 layers, whereas steep connecting ramps are situated in the sandy beds. it is likely that this type of differential beddingparallel thrust faulting and shear brecciation was an important component in translation in the thin-skinned thrust-fault system. it also implies that the stratigraphic succession in a formation that appears to be well preserved may in fact have experienced significant lateral dislocation. footwall synclines the dominant structures in the moserende section are the footwall synclines, which include the growth syncline basins and re-orientated fissure strata. these synsedimentary folds were formed continuously, beginning with the deposition of the sand in a trough. as the thrusting up over the footwall ramp progressed, the trough deepened; contemporaneous with this deepening of the footwall block, the hanging-wall ramp north of the trough started thrusting, which resulted in a drag bend of the northern flank of the syncline. some of the synclines were trapped and overthrust, resulting in overturning of the northern flanks, which in a few cases created nearly isoclinal, recumbent folds. in the piggyback basin of the mr02 thrust sheet, an excellent example of deposition in a growth syncline is preserved. the width of this basin (c. 20 m) is of the same magnitude as the thickness of the sequence deposited and deformed. the basin fill can be divided into four sub-units of the rubjerg knude formation separated by angular discordances (fig. 104). the lowest sub-unit (s 1 ) consists of c. 5 m of trough crossbedded sand deposited in an erosional depression incised into the lønstrup klint formation of the mr02 thrust sheet. it could be argued that in relation to the l/r-unconformity of mr02, this sub-unit shows r-onlap, but it is evident that the trough cross-bedded sand is growth-related to an initial up-thrusting of the mr03 frontal nose. the second sub-unit (s 2 ) was deposited after the first sub-unit was tilted during ramping of the mr02 thrust sheet. the tilting also lifted the s 1 sub-unit up to a level at which the 5 m thick sand package could be subjected to ground-frost. this is deduced from the occurrence of fissure strata emplaced discordantly into the s 1 sub-unit. the fissure strata transecting sub-unit s 1 show small-scale current ripples, which demonstrates that these fissure strata were deposited parallel to the sides of the wedge (fig. 105). the fissure strata can be traced into the second sub-unit s2, which consists of stratified sands characterised by climbing ripple cross-lamination, deposited in a growth syncline trough. the thickness of the s2 sub-unit is 5– 10 m; this sub-unit is dominated by r-onlap in relation to the l/r-unconformity in mr02. deposition of the third sub-unit (s3) was first initiated after sub-units s1 and s2 were deformed by compressional deformation due to push from hangingwall thrusting of mr03 over the footwall block of mr02. the fold structure may be described as an inclined s, where the lower bend of the s corresponds to a footwall syncline. the s3 sub-unit is c. 10 m thick. the base of s3 is an angular discordance on the folded s1 and s2 sub-units. the s3 sub-unit shows r-onlap onto the l/r-unconformity of mr02 and the uppermost part of this sand package covers the detachment anticline structure of mr02 as well as filling the depression on the northern flank of the structure. finally the s3 sub-unit was folded into a footwall syncline by thrust propagation of mr03. in the uppermost part of the basin, a small growth syncline comprises the uppermost sub-unit (s4). this unit is 1–6 m thick and may be regarded as recording deposition in a depression formed by a footwall synclinal bend during the general northward tilting of mr02. interpretation of structural development in the model for the structural development of the moserende section, a distal (southern), an intermediate and a proximal (northern) zone are distinguished. the distal zone includes thrust sheets mr01–mr03, which were probably displaced in similar mode. the intermediate zone includes thrust sheets mr04–mr08, and finally in the proximal zone thrust, sheets mr09– mr13 probably moved sequentially in one continuous displacement. the distal mr01–mr03 zone is regarded initially to have formed one coherent thrust sheet, which was split up during the ramping of mr01. the main bend during ramping took place in mr02, while the trailing end of mr03 still rested on the décollement surface without being elevated, and thus represents the root zone of the mr01–mr03 segment. the sequential development of deposition and deformation is very well illustrated in this first segment with the key locality in the piggyback basin of mr02. here four depositional 136 phases (s1–s4) separated by four deformational phases (f1–f4) have been identified. depositional phase s1. the s1 depositional phase took place during the initial thrusting. about 5 m of trough cross-bedded sand was deposited before (or coeval with) the first deformation phase f1. deformational phase f1 . the first fold phase culminated with the creation of the hanging-wall anticline or detachment anticline in mr02. this folding must be the effect of ramping up from the 20 m level (or 30 m) to the 10 m flat level, resulting in folding of the c. 10 m upper lønstrup klint formation of mr02. depositional phase s2 . during s2 deposition, the angle between bedding in the rubjerg knude formation and the l/r-unconformity is nearly perpendicular. the ramp is not indicated in the ramp cross-section because it was destroyed by mud-mobilisation in the lower part of mr02 (plate 2). deposition of s2 also took place during the exposure of the s1 unit to ground frost conditions as indicated by the presence of fissure strata. deformational phase f2. the f2 folding was related to the hanging-wall thrusting of mr03. as this phase includes thrust displacement of the s1 sand and a synclinal drag of the s 2 sand below the footwall ramp of mr02, it progressed during s 2 deposition. the f 2 phase terminated with the final folding and minor thrust truncation of s 2 . thus the separation of the imbricate thrust sheet mr02 and mr03 took place during the f 2 phase. depositional phase s 3 . deposition of s 3 covered the hanging-wall anticline of mr02, and part of s 2 is discordantly inclined relative to s 3 bedding. s 3 sedimentation is characterised by f-bedding, and displacement probably took place along the 30 m level on top of mr01u, and also st09–st10u. deformational phase f 3 . the f 3 fold phase is restricted to folding of a footwall syncline of the s 3 unit related to the hanging-wall ramp propagation of mr03. it might be interpreted as a continuation of f2, but the mr03 thrusting definitely propagated after s 3 deposition and its extension most likely forms a hangingwall flat over the sand covering the hanging-wall anticline of mr02. depositional phase s 4 . the uppermost growth syncline in the piggyback basin of mr02 was formed by the 5 m thick uppermost unit of the rubjerg knude formation sand. it probably truncated the f3 thrust, and is thus not interpreted to represent dislocated parts of s 1 or s 2 . deformational phase f 4 . the last fold phase created the footwall syncline of the s4 deposits. the hanging-wall ramp of mr03 increased its inclination by about 30° and propagated along a steeper satellite thrust. the footwall syncline was initially overturned to the south, but during the final phase of ramping and steep tilting the syncline became re-orientated into an upright position. the mr04–mr08 intermediate zone initially formed one coherent thrust sheet, comparable to the mr01– mr03 thrust sheets, with a frontal steep ramp, hanging-wall ramp of mr04, and a mr08 trailing-end sheet with the l/r-unconformity below sea level, indicating that this thrust sheet rooted in the décollement zone. in the proximal zone, the mr09 thrust sheet probably propagated as an individual sheet onto the footwall ramps. the main thrusting was initiated with the mr11–mr13 sheets being thrust along a flat in the 30 m level, which is on the lower footwall flat of the trailing end of the mr10 thrust sheet. thus mr11– mr13 can be viewed as thrust-sheet imbrications peeling off the back of mr10. at the latest stage of thrusting, the continued displacement along the décollement zone was responsible for the steepening up of the sheets, somewhat similar to the development of the grønne rende section. for the calculation of the displacement of the individual thrust sheets, the erosionally removed frontal parts have been reconstructed from simple angular geometry, in the same way as the calculation of displacements in the grønne rende section (fig. 11). when the elevation of the l/r-unconformity is considered, it mainly refers to the position of the lowest part of the l/r surface. ideally this part should be horizontal, to indicate the main elevation of the thrust sheet. however, this is not always the situation, and therefore the position of the l/r-unconformity in general refers to its position where a hanging-wall ramp or flat truncates it. mårup kirke section one of the locations where landsliding at present is most dramatic is at mårup kirke (the old church at mårup). this attracts much public attention, not least among the local people. the back-stepping of the head of the slides has been very rapid during the last ten years, and it is probable that the coastal protection measures instituted at lønstrup have increased the erosion of the clif f at mårup kirke. over a number of 137 years the landslides have also obscured the exposures at this location and a lot of interpretation has been necessary to reconstruct the structural framework and the development of the thrusting. the younger yoldia clay and saxicava sand (the vendsyssel formation) cap the mårup kirke section. the erosional unconformity at the base of the vendsyssel formation acts as a drainage surface, which contributes to the generally poor exposure conditions. however, with the experience gained from the other sections, combined with theoretical structural analysis, it is possible to present a model of the structures in the mårup kirke section. there is a marked correlation between the occurrences of piggyback basins comprising the sand-rich rubjerg knude formation, and the build-up of aeolian dunes above the cliff. the northern boundary of the dune field is situated about 200 m south of mårup kirke in the frontal part of the mårup kirke section. south of this area, the aeolian dunes form features ranging from a few metres in height to nearly 50 m high dunes above the rubjerg fyr section, which is also the highest point of the cliff section. thus it is evident that the dunes are formed where a sand source (the rubjerg knude formation) is available (pedersen 1986b). above the main part of the mårup kirke section, no dunes are present, as this is a cliff section that consists only of clay and mud (fig. 108). the most interesting feature at mårup kirke is the packing of the thrust sheets into uniformly developed thrust-fault deformed duplex segments. a theoretical model is presented for the deformation geometry of these duplex units, subjected to extreme compressional development; comparisons with the observations refig. 108. the mårup kirke situated at the head of the cliff in the central part of the mårup kirke section. note that the flat cliff-top surface, representing the horizontal bedding of the vendsyssel formation, is not covered by dunes, reflecting the fact that the clif f consists of clay and mud. in the distance, the rubjerg knude fyr (lighthouse) is being engulfed by sand dunes derived from the sand-rich rubjerg knude formation in the piggyback basins present in the sections to the south. photograph: july 1994; note that by 2002 cliff erosion had reached the corner of the graveyard. 138 corded in the clif f section indicate a reasonable match between the theoretical model and cliff observations. tectonic architecture due to the variation in the distribution of piggyback basins and the general tectonic architecture of compressional framework, the mårup kirke section is divided into three zones: (1) a leading zone (mk01– mk07), (2) a transitional or intermediate zone (mk08– mk10), and (3) a trailing zone, including thrust-fault duplex units annotated mk11–mk20. in the leading zone, the first six thrust sheets have preserved a relatively highly elevated remnant of their strongly eroded piggyback basins. in the transitional zone, only the thrust sheet mk10, just below the mårup kirke, contains a piggyback basin. north of this thrust sheet, the thrust-fault duplex units only comprise the lønstrup klint and stortorn formations. in the intermediate and trailing zones (mk08–mk20), the thrust sheets are defined as duplex units, each comprising four duplex segments, which are annotated d1–d4. thus the lower duplex segment in unit mk09 is annotated mk09d1 and the upper thrust-sheet segment is annotated mk09d4 (plate 2). the southern boundary of the mårup kirke section is the leading-edge thrust fault corresponding to the trailing footwall thrust of mr13 and the hanging-wall ramp of mk01. the boundary is fairly obvious, as it separates the large piggyback basin of mr13 from the c. 50 m thick mk01 thrust sheet. the trailing end of the section is more loosely defined, due to poor exposure and the increasing mud mobilisation of the thrust sheets. it has therefore been defined in relation to the unconformities above the thrust sheets. thus the boundary between the mårup kirke section and the ribjerg section to the north is placed where the fig. 109. upright thrust sheets in the frontal zone of the mårup kirke section (thrust fault arrowed). the dif ference in lithology between the upper and lower levels of the lønstrup klint formation is illustrated by comparing the sand-rich (upper lønstrup klint formation) footwall block (to the right, south) with the mud-dominated (lower lønstrup klint formation) hanging-wall block (left). encircled rucksack for scale. photograph: august 1996. 139 unconformity at the base of the vendsyssel formation truncates the unconformity between the glaciotectonic unit and the ribjerg formation. the glaciotectonic unit is viewed as the unit of deformed deposits related to a glaciotectonic event (pedersen 1993), here included in the rubjerg knude glaciotectonic complex. the defined boundary is very close to where 45° dipping thrust-fault features are obscured both by the mud mobilisation and by superimposed, more or less horizontal, anastomosing jointing. mk01 thrust sheet the first thrust sheet in the leading zone of the mårup kirke section is the nearly 50 m thick mk01 thrust sheet. in the exposed parts of the thrust sheet, the lønstrup klint formation has a thickness of 40 m, which indicates that the deepest level of the section including the lower décollement zone is brought up to the upper flat. the displacement is estimated to be 102 m; this includes construction of the eroded tip of the thrust sheet (fig. 11). this corresponds well with calculation of the displacement according to the equation d × sinα = h, where d is the displacement, α is the angle of the thrust ramping from the lower flat to the upper flat, and h is the distance between the two flats. the distance between the lower and upper flats is 40 m, and assuming an initial thrusting angle of about 23°, the calculated displacement d is c. 100 m. most of the lønstrup klint formation of mk01 is mobilised and forms a diapir-like structure. bedding is only well preserved in the upper 7–10 m of the lønstrup klint formation in this thrust sheet. the bedding is characterised here by medium-bedded sandy turbidites. the rubjerg knude formation of mk01 comprises a nearly 10 m thick succession of glaciofluvial sand. the sand layers show r-onlap and are strongly deformed by a footwall syncline. the l/r-unconformity is situated about 15 m above sea level, indicating an elevation of the lower hanging-wall flat up to the 20 m level footwall flat. this corresponds well with the balanced model for the moserende section, where the trailing-end lower segments mr10u and mr13u still remain to be calculated for. it is thus evident that the mk01 thrust sheet was translated along the 20 m level flat. mk02–mk04 thrust sheets the mk02–mk04 thrust sheets are three relatively small sheets with only minor displacements, about 30 m for mk02 and mk03, and c. 40 m for mk04. the estimates of the displacement for thrust sheets with erosionally removed tips are based on reconstructions following the same principles as applied in the grønne rende section (fig. 11). the thickness of the lønstrup klint formation in the thrust sheets is only between 10 and 20 m, and the thrust sheets are considered to represent imbricates from the upper part of the mk01 thrust sheet, which roots down to the décollement zone in the 40 m flat level. the lønstrup klint formation in mk02 is completely mobilised apart from the uppermost 1–2 m. the mud diapir in mk02 is regarded as an extension of the large mud diapir in mk01. in mk03 and mk04, bedding is partly preserved, and may be compared with the upper sandy bedding seen in mk01. the l/r-unconformity cuts down to at least 2–3 m below the mean level of bedding on the back of mk02 and mk03, and given this uncertainty in location, the position of the unconformity is at the same elevation as in the mk01 and mk04 thrust sheets. it is therefore inferred that the imbricate thrusting of mk02–mk04 was initiated at an early stage, prior to the subsequent deeper-level up-thrusting of mk01. early development of the imbrication is further supported by the thickness of the rubjerg knude formation, which in mk01–mk04 is less than 10 m. the piggyback basin on the back of the non-imbricated mk01 was short-lived relative to those described above in the moserende section where the rubjerg knude formation is up to 25 m thick. along the footwall ramps, sand of the rubjerg knude formation was deposited in growth synclines indicating syntectonic development of the small piggyback basins in the leading zone of the mårup kirke section. mk05–mk07 thrust sheets the mk05–mk07 thrust sheets form an imbricate set of upright thrust sheets, now dipping more than 60°n (fig. 109). the thickness of the thrust sheets is about 20 m in the cliff section. the displacement is 40 m for mk05, 28 m for mk06 and only 15 m for mk07. the lønstrup klint formation, making up most of the thrust sheets, is characterised by a few relatively thick finegrained sand turbidites interbedded with dark blue140 grey silty mud. the bedding is strongly disturbed by thrusting and jointing, similar to the thrust-fault framework described in mr03, and smallto medium-scale duplexes are common. in the upper half of the mk07 thrust sheet, a major footwall syncline is developed below the hanging-wall ramp of mk08. the rubjerg knude formation is only represented in mk06, where the l/r-unconformity is elevated up to about 20 m above sea level. in the mk04–mk05 thrust sheets the elevations are more than 20 m, indicating that these thrust sheets were lifted up and translated along an upper flat resting on three lower segments. the accumulation of these segments forms a subsurface duplex, probably deformed into an intense network of anastomosing thrust faults. the duplex segments constitute the trailing lower segments of mk01–mk04, and for mk06 and mk07 the trailing lower segment of mk05 is also added (see balanced crosssection, plate 2). mk08–mk10 thrust sheets mk08 is the southernmost thrust sheet in the transitional zone of the mårup kirke section. the transitional zone is characterised by the change from thick continuous successions of lithologies into sheet units comprising duplex segments bounded by footwall and hanging-wall flats. along these flats, lateral translation preceded tilting and steepening during propagation along the ramps. the thrust fault separating mk07 and mk08 is the trailing-end footwall ramp on top of mk07 and the hanging-wall flat of mk08. at the top, mk08 is bounded by the footwall flat and hanging-wall ramp between mk08 and mk09. the thrust fault and the general bedding in mk08 dips at 35°n. in the exposed cliff section, mk08 is 40 m thick, which implies that the hanging-wall flat is a segment of the lower décollement zone. above this, four segments may be identified corresponding to the four main levels of flats below the l/r-unconformity. each flat-segment is about 10 m. the displacement along the hanging-wall thrust is c. 110 m, calculated from the equation d × sinα = h, given that the height h is c. 65 m (sum of cliff section and distance down to the décollement surface) and α is 35°. the estimate is based on the assumption that the tip of the hanging-wall ramp in the lowest segment (initially located at the 30 m flat level) only reaches up to the hinge of the footwall ramp at the top of the cliff. in the mk08 thrust sheet, there is no record of the rubjerg knude formation. due to the intense development of landslides below the mårup kirke, the mk09 thrust sheet is very poorly exposed. the interpretation here is based on the space relationships and the structures exposed in mk08 and mk10. from mk08 it is known that the frontal hanging-wall ramp of mk09 dips at 35°. the bedding is more or less horizontal in the cliff section, judging from the occasional features that can be picked out from the photo-geological interpretation. the distance between the leading and trailing thrust fault is about 70–73 m, and the displacement must therefore be about 50–55 m. the interpretation of the mk09 framework is that the upper segment forms a type 3 fault-bend-fold structure, where the sub-segment resting on the upper flat has been eroded away and truncated by the mk10 hanging-wall flat, the two segments in the intermediate levels form s-shaped type 2 structures, and the lower segment forms a type 1 structure with the trailing end of the segment resting on the lower flat (compare with the model in fig. 103). the mk10 thrust sheet is situated below the mårup kirke, and preserves the most proximal piggyback basin containing the rubjerg knude formation. the thickness of the rubjerg knude formation is about 10 m and the basin is deformed into a recumbent footwall syncline. the l/r-unconformity is situated about 10 m a.s.l., indicating elevation above two lower duplex segments in the subsurface. the increase in the dip of the l/r-unconformity and bedding in the lønstrup klint formation from 20° in the frontal part to 35° in the rear part is interpreted as a bend by the hanging-wall ramp propagating over irregularities in the trailing part of mk09. thus part of the subsurface structure must include the geometric adjustments of a splint appearing due to the ramp angle change (see plate 2). this is reflected in the occurrence of a hanging-wall anticline in the trailing end of mk10. mk11–mk20 thrust sheets in the trailing zone of the mårup kirke section, there are no occurrences of the rubjerg knude formation, and the l/r-unconformity has not been identified. the thrust faults are mainly steeply dipping, about 45°, and at the top of the cliff section the thrust sheets are shear-dragged southwards and reworked into glacitectonites, a truncated glaciotectonic unconformity and local till. 141 this trailing zone can be divided into ten fairly uniform thrust-fault duplex units, each about 55 m long (measured horizontally along the beach level) and separated by 45° steeply dipping thrust faults. for the characterisation and structural explanation of these thrust-fault duplex units, a thrust-ramp-propagation model is presented below (see fig. 114). sedimentary units in the mårup kirke section, no significant additional data have been obtained to supplement the sedimentological descriptions. however, it is evident that the rubjerg knude formation only occurs in the southern part of the section, where it is less than 10 m thick. it is inferred, therefore, that the piggyback basin in the mårup kirke section was short-lived relative to the thicker successions of the rubjerg knude formation further south. the rubjerg knude formation may never have been deposited in the trailing end of the mårup kirke section. structures a conspicuous feature of the mårup kirke section is the shear drag at the top of the cliff section. the main structures formed during the subglacial drag are southerly overturned to recumbent synclines that developed in a sandy glacitectonite about 1 m in thickness. the glacitectonite is interpreted to have formed by subglacial shear deformation superimposed on the proglacially formed thrust-fault and duplex structures (pedersen 1988, 1996, 2000). over a large part of the mårup kirke section, the amount of mobilisation is not very high. this permits a characterisation and interpretation of the deformation of duplex segments and stacking of duplex units, as described below. interpretation of structural development the main purpose of this section is to present an analytical structural model that can be used to interpret the thrust-fault framework developed in the mårup kirke section. the basic elements of the model are the duplex segments, and the structural deformation can be characterised as fault-bend folding. the result of the deformation is a compressional stacking of duplex segments into duplex units with a certain geometry and size. the interpretation of these structures adds to the basis for the discussion of structural developments that concludes this section. fault-bend-fold model for duplex units it was demonstrated in fig. 103 how a duplex unit developed with type 1–3 duplex-segment structures. however, a number of structural configurations may develop from the deformation of duplex segments stacked into duplex units, depending on the amount of displacement and the initial length of the thrust sheet. for the interpretation of the structures in the northern part of the mårup kirke section, as well as a major part of the ribjerg section, the analytical models in fig. 110 have been constructed. the premises for the models are: (1) the duplex unit comprises four initially horizontal sheets with a thickness of 10 m each, (2) the bounding leading and trailing thrust ramps dip at 45° (maximum angle of thrust-fracture formation), (3) the vertical distance between the lower and upper flat is 40 m, and (4) the lateral compression cannot exceed the packing of the sheets in 45° dipping imbricates. from the latter premise, it can be predicted by simple trigonometric calculation that the lateral distance between the bounding thrusts of the duplexes should be close to 56.5 m, and this corresponds very well with the thrust features recorded in the cliff section. to illustrate the model, one ideal case is considered, namely the case where the displacement is to the top of the ramp, which has the same length as the maximum compression distance of 56.5 m (fig. 110, type 3). the displacement takes place along the lower footwall flat, and the lower hanging-wall flat propagates up along the 45° dipping footwall ramp. the resulting structural framework is an l-type fault-bend folding (fig. 103). in this case, the lower thrust segment will just reach the level of the upper flat. the displacement is c. 42 m and the balanced length of the duplex unit is 98.5 m, which results in a calculated compression of 43%. due to the propagation along the upper flat and the ramp-bend folding, a hangingwall anticline is formed, which can be described as a fairly upright, angular antiformal stack. the case described above is shown as type 3 in fig. 110. this case might also be called the angular antiformal stack type. further cases can be considered with decreasing or increasing displacement relative 142 l0 (m) 70.5 1 20%compression 2 33% 3 43% 4 50% 5 56% 6 60% 84.5 98.5 113 126 140 7 l1 56.5 comp. = 28 84.5 33% l0δ 1 l1 56.5 comp. = 14 70,5 20% l0δ 2 l1 56.5 comp. = 28 84,5 33% l0δ 3 l1 56.5 comp. = 42 98.5 43% l0δ 4 l1 56.5 comp. = 56.5 113 50% l0δ 5 l1 56.5 comp. = 70.5 127 56% l0δ 6 l1 56.5 comp. = 84 140.5 60% l0δ 8 l1 56.5 comp. = 56.5 113 50% l0δ 9 l1 56.5 comp. = 70 126.5 56% l0δ fig. 110. the duplex-unit model for fault-bend folding of duplex segments. the basic elements for the constructed models are: (1) the duplex unit comprises four initially horizontal sheets, separated by thrust-fault flats, and each sheet is 10 m thick, (2) the bounding leading and trailing thrust ramps dip at 45° (maximum angle of thrust-fracture formation), (3) the vertical distance between the lower and upper flat is 40 m in types 1–6, and in types 7–9 it is extended 10 and 20 m above the upper footwall hinge, and (4) the compression cannot exceed the packing of the sheets in 45° dipping imbricates. the cases in the model are selected with steps jumping one 10 m level from case to case. according to simple trigonometry, this will result in a displacement unit of c. 14 m, and multiples of this. the initial dimensions (l 0 ) of duplex units 1–6 are given by the scale in the lower right corner. type 1 is a single monoclinic flexural kink fold. type 2 is a double monoclinic flexural kink fold. type 3 is an angular antiformal stack. type 4 is a flat-topped antiformal stack. type 5 is a lateral extension of the flat-topped antiformal stack. type 6 is a lateral extension of the flat-topped antiformal stack, where it is demonstrated that the frontal limb in the antiformal stack retains its profile, and it is only a lateral translation of duplex segments that responds to the further compression of the duplex unit. type 7 is a per fect g-s-l structure (fig. 107). compression of 33% in type 7, and an elevation of the ramp by two 10 m levels, results in the monoclinic flexural kink fold. the ef fect of increasing compression and propagation up along the extended ramp (types 8 and 9), demonstrates the development of the antiformal stack in a manner comparable to that from type 3 to type 4. note that the given maximum stacking of imbricates constrains the size of the balanced length of duplex units. this is demonstrated in the diagram relating balanced length to magnitude of compression. l1, length after deformation; δ, shortening; lo, initial length; comp., compression. to type 3. if the displacement of the lower thrust-sheet segment is less than the height of the footwall ramp, the duplex segments above have two ramps to pass. consequently, two ramp-bend folds will be created, which may also be described as a repeated monoclinic flexural kink-fold (fig. 110, type 2). in the model, the second ramp-bend fold will not develop until the displacement exceeds 20%, corresponding to a displacement lift of only one 10 m level (fig. 110, type 1). with increased compression, the hanging-wall anticline formed in type 3 will develop into a flat-topped antiformal stack (types 4 and 5). finally, type 6 dem143 onstrates the lateral extension of the flat-topped antiformal stack resulting from 60% compression. note that in this case the frontal limb in the antiformal stack maintains its profile and an increase in compression only results in lateral translation of thrust sheets. increasing compression also requires increasing length (l0) of the duplex unit, which is demonstrated by the diagram in the lower right corner of fig. 110. in fig. 110, the cases with an extended ramp have also been examined. this corresponds to thrusting above the upper hinge of the footwall ramp, which would be the case if syntectonic sedimentary units were deposited on the upper flat preceding ramp propagation. type 7 is a perfect γ -s-l structure, exemplifying this development. it is formed by compression of 33% and elevation of the ramp by two 10 m levels, here creating a monoclinic flexural kink-folding. the effect of increasing compression and elevation of the ramp from type 8 to type 9 demonstrates the development of the antiformal stack in a manner rather similar to that from type 3 to type 4. characterisation of thrust duplex mk11–mk20 on the basis of the models in fig. 110, the mk11 thrust sheet is classified as a type 2 structure due to the presence of two monoclinal flexures. however, the structure in mk11 must incorporate the effects of the displacement of the thrust segment of mk10. consequently, the upper segments of mk11 are stacked on each other as relatively short duplex segments. moreover, the topmost part of mk11 is dragged out and sheared over the piggyback basin of mk10. this dragged part can be interpreted as the frontal limb of the antiformal stack initially formed over the upper footwall hinge. mk12 is the duplex unit situated north of mårup kirke. all the structures dip at 45°, except for the uppermost shear-dragged parts, which were reworked into a local till (the kattegat till formation). thus the structure is interpreted mainly as an l-structure, probably a type 5 or 8 structure with 55% compression and a balanced length of c. 126 m (fig. 110). mk13 has an undulating flat-lying structure with flexural drag up along the footwall ramp. it is thus interpreted as a type 4 structure with a flat-topped antiformal stack capping the frontal part of the lower thrust segment, which was only displaced up to the reference level of the l/r-unconformity. compression amounts to 50–55%, and the balanced length is estimated to be 120 m. mk14 is considered to be similar to mk13. it was probably very close to the modelled type 4 structure (fig. 110), prior to glaciotectonic shearing and truncation of its flat-topped antiformal stack. mk15 and mk16 are probably the closest approximation to a perfect γ-s-l-structure of type 7 in the model (fig. 110). mk17 and mk18 may well be inferred to be of the same type. however, the exposures are here too poor for definitive structural characterisation. in mk19 and mk20, structures with 45° steep dips are displayed in the cliff section. these duplex units can thus be interpreted as type 9 duplexes. discussion of structural development although the balanced section is subject to some uncertainties in the mårup kirke section, calculation of the compression from the measured length of the section l1 = 978 m, and a balanced length of about l0 = 1814 m gives 46%. as described above, the section is divided into three architectural zones: (1) a leading zone (mk1–mk7), (2) a transitional or intermediate zone (mk8–mk10), and (3) a trailing zone that includes thrust-fault duplex units (mk11–mk20). the discussion below attempts to demonstrate the proximal–distal thrust-fault development. the fault-bend-fold model for duplex units describes the thrust-fault structures in the trailing zone and gives an approximation of the structural framework of the major part of the mårup kirke section. the absence of the rubjerg knude formation in the trailing zone suggests that it was never deposited here. moreover, the thrust stacking of the duplex units started before, or just at the beginning of, deposition of the rubjerg knude formation in the most proximal part of the glaciotectonic complex. it is further suggested that the thrust levels rapidly shifted to lower levels in progressive steps. so, after the first few hundred metres of peeling off the uppermost thrust segments, the thrusting propagated for the next five hundred metres in the intermediate flat levels. finally, the main compression started to stack the duplex units up into imbricates during translation along the lower flat level, the décollement zone, and differential displacement between the duplex segments. the displacements of the duplex units were limited by the maximum shortening between the 45° steep northward-dipping ramps. it might be suggested that the displacement was much larger and considerable amounts of the leading part of the thrust sheets were 144 eroded away from the upper flat. however, this is unlikely for two reasons: (1) the amount of compression is 50–60% which is considered to be a limiting amount of compression for natural systems, and (2) the structures discernible from the photo-geologically interpreted cross-section support a model with c. 50% shortening. another suggestion could be that the duplex imbricates were formed subglacially, bounded by a floor thrust (the décollement zone) and a roof thrust situated in the glaciotectonic unconformity (the sole of the glacier). this is disproved by the fact that the antiformal stack above the duplex units would have required space to be stacked up on the upper flat. thus, although the antiformal stacks were removed by glacial erosion and the upper part of the mårup kirke section is shear-dragged and truncated by the glaciotectonic unconformity (formed subglacially), the thin-skinned thrust faulting developed in a proglacial setting in front of a progressively advancing ice margin. mk08 is the leading thrust sheet in the transitional zone of the mårup kirke section. it has a considerable displacement, more than 100 m, and it probably ramped up to the 20 m flat level along which it was translated for more than 50 m before its hanging-wall flat propagated up to the uppermost footwall flat. thus, all the segments in the thrust sheet were earlier translated along the various flats before mk08 was displaced up along the footwall ramp on mk07. there should therefore be a stepping down of the trailingend sheet in the zone. this would correspond to translation along the lower flat level of the mk09–mk10 thrust sheets, which facilitated the formation of a depression above mk10, where the rubjerg knude formation was deposited and preserved in the most proximal piggyback basin of the glaciotectonic complex. the leading zone is characterised by carrying a relatively high-elevated piggyback basin, where the rubjerg knude formation was deposited on an uneven erosional unconformity. during the early phase of imbrication, this piggyback basin was separated into five sub-basins, before they were finally trapped by overthrusting and deposition ceased. the accumulated displacement in the leading zone is c. 180 m. the thrusting probably started from a detachment level in the upper flat (10 m level), inferred from the thickness of mk02–mk04. thrusting then shifted down to the second flat (20 m level). assuming that the first half of the displacement started as an imbrication of the mk02–mk06 thrust sheets, then lateral translation of the upper 10 m thrust-sheet segment resulted in lateral displacement of the upper thrust level in the order of 100 m. subsequently, the detachment surface was lowered down to the 20 m level, and it is evident that this detachment surface is the next flat level, along which about 100 m lateral translation occurred. one of the main lines of evidence that this level is another pervasive flat level is that it acted as an upper flat for the displacement of mk01. the propagation of this thrust sheet probably started with a minor dislocation along the 30 m flat level, which is known to be a pervasive flat level from the moserende section, before it moved down to be a dislocation along the lower décollement level (40 m flat level). from the lower décollement level, mk01 ramped up to the 20 m flat level along which translation occurred over a distance of 80 m before its hanging-wall flat and ramp was ramped up to the surface along the footwall ramp at the trailing ramp of the moserende section. during displacement, the mk01 thrust sheet carried the mk02–mk07 sheets piggyback resulting in over-steepening of these thrust sheets towards the trailing end (mk05 and mk07). ribjerg section the northern termination of the rubjerg knude glaciotectonic complex is the sandy hill at lønstrup called ribjerg. most of the coastal clif f below ribjerg is now protected, and vegetation covers the cliff exposures at ribjerg. however, on the south-western side of ribjerg a funnel-shaped gully has been formed by steady erosion due to high groundwater drainage in the glaciofluvial sand (fig. 111). at the boundary between the sand and the underlying mud, groundwater wells up and creates quicksand. thus, although a section through the glaciofluvial sand is well exposed, access is difficult and potentially dangerous. in spite of such obstacles, a detailed log of the succession has been measured, and the locality yields the type section of the ribjerg formation. in addition, the section is the site for studying the glaciotectonic unconformity above the skærumhede group, cropping out at the ‘lille blå’ (northernmost part of the cross-section in plate 1). ‘store blå’ and ‘lille blå’ north of the mårup kirke section, the unconformity above the mud-rich lønstrup klint formation dips gently to the north. jessen (1918, 1931) named this part of the cliff ‘det store blå’ and ‘det lille blå’ (the 145 big blue and the small blue, respectively, a reference to the blue colour of the clayey mud in the mud-rich part of the cliff section). in general, the mud is a mobilised succession with only few bedding surfaces and thrust faults preserved. in the store blå cliff section, the structural features recorded accord well with the maximum compressional model described for the duplex units of the mårup kirke section (fig. 110). in the lille blå cliff section, the mud is structureless, and no primary bedding surfaces are preserved. a secondary sub-horizontal planar fabric is recognisable, and pebbles and boulders occur on the unconformity as well as in the uppermost metre just below the unconformity. jessen (1931) interpreted the lille blå as dislocated older yoldia clay, which he named portlandia arctica clay after the occurrence of the identified mollusc species in the unit. jessen’s description of the clay compares well with the characterisation of the skærumhede group and the model of glaciodynamic development presented below. due to the progressive deformation in the proximal part of the glaciotectonic complex, deeper levels of the skærumhede group were thrust up into a position close to the main l/r-unconformity level, such that an increasing proportion of the group has been eroded. jessen (1931) also described another important feature related to the lille blå cliff section. before 1895, it could be observed that the unconformity was folded into a syncline with a fold axis directed n–s. this is of course unusual since all structures described until now are assumed to have been formed by compression directed n–s due to the advance of the ice cap from the north, resulting in mainly e–w-trending structural features. the n–s-orientated fold axis is interpreted to be related to deformation by ice advance from the east, an event that also deposited the mid danish till formation. tectonic architecture the ribjerg section is defined as the section between the northern boundary of the mårup kirke section and the end of the lønstrup klint cliff section, which terminates at the vegetation-covered cliffs below the town of lønstrup. the southern boundary of the section is situated where four unconformities are superfig. 111. the ribjerg section viewed towards the north. the sandy cliff in the centre of the figure is the type locality of the ribjerg formation. photograph: july 1994. 146 imposed upon each other. these are: (1) the l/r-unconformity, (2) the glaciotectonic unconformity below the kattegat till formation, (3) the unconformity between the kattegat till formation and the glaciodynamic succession related to the ne-ice advance, and finally (4) the unconformity between the glaciodynamic successions and the vendsyssel formation (plate 1). the first three unconformities are here collectively termed the blå-unconformity. the blå-unconformity dips at 2–3° to the north. the surface is relatively planar, but uneven. a few clasts remain in depressions on the surface, but clasts protruding into the surface from below are more common. the unconformity between the ribjerg formation and the vendsyssel formation is an erosional surface dipping gently to the south. the main lithology in the vendsyssel formation is the saxicava sand, which consists of sandy heteroliths. these beds onlap the unconformity, which probably was subaerially exposed before inundation by the rising younger yoldia sea. sedimentary units in the ribjerg section, four sedimentary units are represented: the skærumhede group, the ribjerg formation, the mid danish till formation and the vendsyssel formation (figs 14, 17, 33). skærumhede group the skærumhede group comprises two formations: the stortorn formation and the lønstrup klint formation. in the southernmost part of the section (at the store blå), it is possible to distinguish the two formations (fig. 17). however, in the northern part of the ribjerg section, pervasive mobilisation has obliterated the primary lithological differences and the sediments may only be referred, undifferentiated, to the skærumhede group. in the southern part of the section, the stortorn formation constitutes the lowermost 10 m of the cliff section (fig. 17). here a cataclastic breccia separates the stortorn formation from the lønstrup klint formation above. it is inferred that this breccia represents one of the thrust-fault flats that form the boundary of the duplex segments building up the duplex units of the section. the skærumhede group is truncated by the blåunconformity, above which the vendsyssel formation was deposited. blå-unconformity the blå-unconformity is considered to represent three superimposed unconformities. the first one is the l/runconformity, the existence of which is only rarely demonstrable in this section. the second unconformity is the glaciotectonic unconformity below the kattegat till formation. the kattegat till formation has been almost completely eroded away from the ribjerg section, but is present in small, isolated pockets (fig. 31). however, the glacitectonite related to the subglacial deformation below the kattegat till formation is well preserved in a zone more than 1 m thick below the blå-unconformity (fig. 32). erratic clasts are common in this zone, probably lodged into the soft sediment from the till above, and an indicator boulder of larvikite has been recognised. a number of clast fabrics have been measured, which show a n–s long-axis orientation (variation from 010° to 175°). the unconformity is preserved at the base of the vendsyssel formation in the northern part of the mårup kirke section (fig. 37). the third unconformity is the erosional surface upon which the ribjerg formation was deposited. the creation of this surface removed much of the evidence of the preceding unconformities; indeed, at the southern extent of the unconformity, the ribjerg formation is also absent, and the vendsyssel formation rests on the composite surface. ribjerg formation the c. 25 m thick glaciofluvial sand of the ribjerg formation dominates the ribjerg section (fig. 33, plate 1). the formation comprises fineto medium-grained sand, coarsening upwards into gravel-dominated beds at the top (fig. 33). the formation was deposited on the erosional unconformity capping the skærumhede group (the blå-unconformity). at this surface, a residual coarse clastic bed is present, less than half a metre in thickness, dominated by clayey clasts derived from the unit below. the clayey clasts continue to appear in the sand beds in the lowermost 5 m of the formation. the middle part of the formation is dominated by trough cross-bedding, and the flow direction indicat147 ed from measurements of foreset beds was from east to west. the fill of the large channels incised into the medium-grained sand package also include gravel and slumped diamictite material. water-escape dykes and sand-filled cracks are common in the sand within the large channels (fig. 34). the formation coarsens upwards into a trough cross-bedded sandy gravel in the uppermost 3 m, just below the diamictite referred to the mid danish till formation. mid danish till formation the mid danish till formation is a c. 3 m thick unit of grey brown to light yellowish brown sandy till that overlies the ribjerg formation (figs 14, 33, 35). the till is divided into lower and upper beds. the lower bed is a laminated to thin-bedded, fine-grained sandy, matrix-supported diamict. lamination and bedding is deformed into irregular intraformational slump folds with fold axes trending n–s, indicating a slump-slide direction towards the west, and the unit is interpreted as a sediment gravity flow or flow till (dreimanis 1988). the upper bed is a massive, structureless and sandy matrix-supported diamict (fig. 35). the clasts, pebble to cobble in size, occur randomly, and the till fabric shows an a-axis orientation gently dipping towards the east. the unit is interpreted as a basal lodgement till (dreimanis 1988) superposed on the flow till and deposited by an ice stream moving from east to west. the mid danish till formation is truncated by the erosional unconformity upon which the vendsyssel formation was deposited. vendsyssel formation in the ribjerg section, the vendsyssel formation truncates the mid danish till formation, the ribjerg formation and the blå-unconformity. the maximum thickness in this part of the lønstrup klint section is about 12 m, decreasing towards the north, where it onlaps the unconformity above the ribjerg and mid danish till formations (fig. 33). the vendsyssel formation comprises laminated mud and thin-bedded finegrained sandy heteroliths, which in the southern part of the ribjerg section are characterised by well-preserved trace fossils created by the bivalve hiatella arctica, often with the shells preserved in life position (fig. 41). structures in the ribjerg section, the most important structures are the anastomosing joints related to the glacitectonite below the blå-unconformity (fig. 32). at the lille blå locality, the rhomb-shaped segments, 0.5–3 m in size, bounded by conjugate shear joints, are flatlying. the angle between conjugate joints varies from 10–30° and the zone-axis is orientated more or less e–w. at the store blå locality, the shear joints are more parallel with a spacing of c. 30 cm between the almost horizontal fractures, and in the southernmost part of the section, sand-fill intruded the fractures to create rhomb-shaped segments in a sandy mud matrix. interpretation of glacial geology and stratigraphic development in the interpretation presented here, the blå-unconformity is considered to be a modulation surface or deformational layer below the advancing front of the norwegian ice. the unconformity may even be interpreted as the surface onto which the sole of the ice pressed during the propagation towards the glaciotectonic complex developing in front of it. after the ice retreated, a hill-and-hole pair formed. rubjerg knude is here viewed as the hill and the depression extending to the north of the northward-dipping unconformity corresponds to the hole. the hole was subsequently filled with glaciofluvial sands (the ribjerg formation) that are younger than the rubjerg knude formation. on top of the ribjerg formation, jessen (1931) described a sandy till that is here referred to the mid danish till formation, but he also recorded a single till-bed intercalated in the meltwater sand. this sandy till as well as the thin diamictite layers related to the slumps in the troughs and channels are interpreted as precursors to the flow till that initiated deposition of the mid danish till formation. the ribjerg and mid danish till formations were formed as proglacial and subglacial units during the advance of the ice from the east towards the west with a source area in central sweden. this ice advance was also responsible for the gentle folding of the blå-unconformity and the beds above it into a syncline with a n–s-trending axis, as noted by jessen (1931). 148 dynamic development of the thin-skinned thrust faulting the dynamic development of the thin-skinned thrust faulting in the rubjerg knude glaciotectonic complex is presented as a sequence of restoration stages. thus, the progressive deformation of thin-skinned thrust faulting and related syntectonic depositional developments are illustrated in sequentially restored crosssections beginning with the proximal moserende section and concluding with the ulstrup section in the most distal part of the thrust-fault complex. the basis of each restoration sequence is the balanced profile (plate 2a), and the end stage is identical with the thrustfault cross-section (plate 2b), including the interpretation of the unexposed ramps and flats in the subsurface. the most proximal sections, the mårup kirke and the ribjerg sections were interpreted individually in the preceding chapters, and are not included here. in a summary scheme (see fig. 123), it is concluded that the dynamic development was a process of continuous progressive deformation. thus, although the following description is concentrated on the individual sections, it should be kept in mind that there is overlap between sections, and that the whole system was mobile. thus a displacement of 5 m on one thrust might be followed by 10 m on a more proximal thrust and 7 m on a more distal thrust depending on the local conditions. this is the reason why a number of displacements appear to be out-of-sequence, but within limits that respect the lowest décollement level, and that displacements along the most distal, leading-edge thrusts were the last to be activated. it is therefore also evident that displacement along a leading-edge ramp may correspond to a translation along a flat in a proximal section. moserende section the thrust-fault development in the moserende section is regarded as normal progressive piggyback thrusting from the proximal towards the distal part. the mr12 thrust sheet was probably the first to be thrust onto the relatively thinner piggyback basin on the back of mr11 after a c. 10 m thickness of rubjerg knude formation sediments had been deposited. this is included in the first stage of the sequential restoration (fig. 112, stage 1). a total of eight stages have been dif ferentiated, of which stages 1–6 are illustrated in fig. 112. the stage preceding the deformation is shown in plate 2a, and the final stage terminating the deformation is reconstructed in plate 2b. moserende stage 1. the initial thrusting started with 40 m displacement of mr12 over the back of what was to become mr11. this thrusting was rooted down to the 20 m intermediate décollement level. during accumulation of a 20 m thick succession of sediments in the piggyback basins above the mr13 and mr12 sheets, the thrusting progressed with ramping of mr11 over the rubjerg knude formation on top of what was to become mr10. this thrusting involved ramping and translation of the lower segments of mr13– mr11 from the 30 m flat level onto the 20 m flat level. the trailing-end segments of the moserende section were contemporaneously over-thrust by mk01, the frontal thrust of the mårup kirke section, which is rooted in the 40 m décollement level. the accumulated displacement of thrusting of mr13, mr12 and mr11 is estimated at about 150 m. moserende stage 2. thrusting of mr09 initiated this stage. the mr09 thrusting ramped up from the 40 m décollement level, and a single duplex formed during stacking of the lower mr09 thrust segment. the mr09 sheet was displaced c. 40 m over the mr08 piggyback basin. contemporaneously, mr10 was thrusted over mr09 and the mr10 hanging-wall flat extended from the top flat level down to the 40 m décollement level. the mr13–mr11 thrust sheets were then passively translated on the trailing lower segment of mr10. moserende stage 3. initial imbrication of the mr08– mr05 thrust sheets resulted in an accumulated displacement of c. 200 m. the ramping was rooted in the 40 m décollement level along which the main translation of the trailing-end thrust sheets of the moserende section took place. the thrusting involved a complex relationship between mr07 and mr06 that may be interpreted as a connecting splay duplex (mitra & sussman 1997). above the l/r-unconformity, the deposits of the rubjerg knude formation probably reached a thickness of 20 m. moserende stage 4. the frontal part of the section was activated by c. 40 m translation of mr1 along the 30 m 149 décollement level over the lowermost trailing-end segments in the stortorn section. mr02 and mr03 followed this translation, whereas mr04 ramped up one level from the 40 m décollement level to the 30 m flat level that resulted in the initial ramping of mr04 up over the rubjerg knude formation on the back of mr03. the continued displacement consequently reorientated thrust sheets mr05–mr07 into more steeply dipping orientations. the trailing-end thrust sheets from mr08 and northwards were translated passively during this displacement. moserende stage 5. the frontal displacement of mr01 continued along the 20 m flat level over the trailingend segment of the stortorn section. mr02 ramped up along the footwall ramp at the trailing end of mr01 during a fault-bend rotation, which also included the lower segment of mr01u. a vertical thrust separation of c. 10 m brought mr02 up along the northern termination of the mr01 piggyback basin. during the passage of two intermediate ramps, an irregular anticline formed on the back of mr02 that had significant implications for the synsedimentary structures formed in the mr02 piggyback basin (see description of the moserende section, above). during mr04 thrusting, mr03 was imbricated along the upper 10 m flat level and the mr03b and mr03c thrust segments started to break through the piggyback basin. from the rear, mr04 was pushed by mr05 which had to pass up over the fault-bend-folded segment mr04u. together with mr06 and mr07, the mr05 thrust sheet moved up to the highest level indicated by the l/r-unconformity, situated c. 20 m above sea level on the back of mr06 and mr07, and their thrust faults were steepened into a nearly vertical position. moserende stage 6. in the frontal part of the moserende section, mr01 picked up a lower segment and thrust up to the 20 m flat level, which consequently also elevated the piggyback basin up into its present high level. the trailing-end ramp of mr01 formed the footwall ramp for the mr02 thrusting, which resulted in a fault bend of mr02 as well as mr03. this was followed by the final displacement of 18 m along the leading mr03 thrust. minor adjustments and re-orientation of mr04–mr07 followed the ramping of mr03, and the trailing-end thrust sheets mr08–mr13 were passively displaced by translation along the 40 m décollement level. moserende stage 7. during this stage, a complex duplex was formed by thrusting of the frontal lower segments, which also including the trailing-end segments of the stortorn section. moserende stage 8. the final displacement along the leading thrust-fault ramp in the moserende section progressed up along the stortorn trailing-end footwall ramp. moreover, the fault-bend folding due to thrusting in the stortorn section brought the thrust sheets into their present steeply dipping orientation. moserende section: summary data balanced length (l0): 1120 m cross-section length (l1): 650 m shortening (δl): 470 m compression: 40.2% stortorn section the most important development in the stortorn section was the change from the lowermost 40 m décollement level to the 30 m décollement level. the ramp, or progressive development of lower ramps, which marked the change, is here referred to as the stortorn lower segment footwall ramp, and was located somewhere near the thrust between st04 and st03. thus, the st03–st01 thrust sheets had their lower décollement level at 30 m, whereas the thrust faults related to st04–st10 were rooted in the 40 m décollement level. formation of a duplex complex comprising the lowermost thrust segments exposed the stortorn formation, the oldest strata involved in the thrusting. in the frontal part of the section, a complex stacking of lower segments, remaining in the subsurface from displacement in the grønne rende section, resulted in duplex formation that elevated st01–st03 about 30 m above the reference level. due to arguments presented later (see grønne rende section) the duplex stacking had to have been contemporaneous with the shortening of the grønne rende section. in the stortorn section, seven stages have been differentiated of which stages 1–5 are illustrated in fig. 113. stortorn stage 1. this stage is a direct continuation of the displacement in moserende stage 4. in the stortorn section, deformation was initiated by imbrication of st07, st09 and st10 with an accumulated displacement of about 50 m. this resulted in a ramping 150 skærumhede group active thrust fault rubjerg knude formation moserende 2 mr13 mr12 mr11 m mk1 moserende 3 mk1 mr13 mr12 mr11 moserende 4 mk1 mr moserende 5 mk1 mr13 moserende 6 mk1100 m n s moserende 1 mr13 mr12 mr11 mr10 mk1 fig. 112. dynamic model of progressive deformation in the moserende section illustrated in six sequential restoration cross-sections. the six stages demonstrate steps in the development between the balanced cross-section and the structural cross-section (plate 2); thus the starting and final positions are not shown. the red lines indicate the active displacement surfaces in each deformation stage. the basic décollement sur face is the 40 m flat level. from this, the flat levels rise by 10 m onto the reference level (l/runconformity) defined as the 0-level. note (1) that the final two stages (7, 8) discussed in the text are not illustrated, and (2) that the thrustsheet terminology in figs 112–122 is simplified (i.e. mr 3 on fig. 112 is equivalent to mr03 in the text). 151 stortorn footwall ramp stortorn footwall ramp stortorn footwall ramp mr10 mr9 mr8 mr7 mr6 mr5 mr4 mr1 st10 mr10 mr9 mr8 mr7 mr6 mr5 mr1 st10 mr1 st10 r13 mr12 mr11 mr10 mr9 mr8 mr7 mr6 mr5 mr4 mr3 mr2 stortorn footwall ramps 3 mr12 mr11 mr10 mr9 mr8 mr7 mr6 mr5 mr4 mr4u mr3a mr3u mr3c mr3b mr2 mr1a mr1u mr1 st10 st10 stortorn footwall ramps mr13 mr12 mr11 mr10 mr7 mr6 mr5 mr4 mr3c mr3b mr3a mr2 mr1 st10 st10u st10u mr8 mr9 stortorn footwall ramp mr9 mr8 st10 mr1 152 100 m n s skærumhede group active thrust fault rubjerg knude formation mr1 st10 st9 st8 st7 st6 stortorn 2 st5 mr1 mr1 st10u st10 st9 st8 st7 st6 st5ust8ust8ust9u stortorn 1 stortorn footwall ramp st10mr1 st9 st6 st7 st8 stortorn 4 stortorn footwall ramps stortorn 5 stortorn footwall ramps st10 mr1 st10 st9 st8 stortorn 3 stortorn footwall ramps fig. 113. dynamic model of progressive deformation in the stortorn section illustrated in five sequential restoration cross-sections. the cross-sections demonstrate steps in the development between the balanced crosssection and the structural cross-section (plate 2); thus the starting and final positions are not shown. the red lines indicate the active displacement sur faces in each deformation stage. note that the duplex segments gr u refer to elements that had to be deformed contemporaneously with the shortening taking place in the grønne rende section; duplex segments st s refer to horse/splint segments. the final two stages (6, 7) discussed in the text are not illustrated. 153 st4 rf6 st1s st1st2 st3 st4 st4s st4s st4s st4s st2s st1s stortorn lower ramp rubjerg knude fyr ramp frontal footwall ramp st5 st5 st4 st2st3 rubjerg knude fyr footwall ramp rubjerg knude fyr footwall ramp stortorn lower segment footwall ramps rubjerg knude fyr footwall ramp stortorn lower segment footwall ramps st9 st7 st6 st5 st4 st3 st2 st1 rf6gru gru gru st8 st6 st5 st4 st3 st2 st1 st2s rf6 stortorn lower footwall ramp rf6 st1s st1 st1s rf6 154 up of st10 from décollement level 40 m to flat level 30 m, along which the translation displacement took place. most of the st09 thrust sheet was also ramped up by the formation of a lower duplex structure. both st10 and st09 were affected by fault-bend folding, which created a major distortion of the l/r-unconformity surface in the uppermost part of the thrust sheets. stortorn stage 2. during an accumulated displacement of about 200 m related to the st10 and st09 thrusts, ramping progressed with development of the first imbrications of st08 and st06. due to ramping from the lowest décollement level to the 30 m flat level in the trailing end of st06, a fault bend affected the st07– st10 thrust sheets that were translated piggyback on the st06 thrust sheet. this contributed to the steepening up of the st07–st10 thrust structures. in the frontal part of the section, the imbricate thrusting was initiated at st01–st03. accumulation of sediments referred to the rubjerg knude formation reached a maximum thickness of about 20 m, notably in the synformal troughs of st03 and st09 that formed during the progress over the ramps below. stortorn stage 3. at this stage, st05 was thrust about 40 m up over the footwall ramp on the back of st04. the st05 thrust was rooted in the 30 m flat level, and during a passage of a lower ramp from flat level 30 m to 20 m, the initial fault-bend-fold resulted in undulation of the l/r-unconformity at the top of the st05 thrust sheet. the st06 thrust sheet progressed over the footwall flat of st05, and both thrust faults were rooted down to the 30 m flat level along which the main translation of the sheets emplaced piggyback on st06 took place. the st08 thrust sheet was finally displaced along the upper flat at the top of the st07 piggyback basin. consequently, most of the 20 m thick succession in this piggyback basin was preserved and indicates the maximum level of sediment accumulation in the rubjerg knude formation during stage 3. thrusting of the st08 sheet along the footwall ramp on the back of st07 resulted in a further steepening of st09 and st10, while the frontal elevated parts of the st08–st09 thrust sheets became subject to erosion. the trailing-end lower segments of st10–st07 were over-thrust by the frontal parts of mr01 and mr02, corresponding to stage 7 in the moserende section. the accumulated displacement in stortorn stage 3 was of the order of 320 m. stortorn stage 4. during this stage, the st04 thrust sheet was thrust 40 m over the piggyback basin of st03, and st05 was thrust about 70 m over the upper flat on top of the piggyback basin of st04. during this relatively large displacement of st05, two lower duplex segments were picked up from the lower 40 m décollement level. after ramping over the stortorn lower ramp, the duplex segments participated in the thrusting up along the footwall ramp on the back of st04. the lower trailing-end segments of the stortorn section were finally thrust up along the steep footwall ramp on the back of st10 and subsequently the frontal parts of the moserende section were brought into their present upright orientation. erosion and redeposition affected the piggyback basins on st05 and st08, whereas thrusting over st07 and st04 sealed these piggyback basins. the accumulated displacement reached about 410 m. stortorn stage 5. a substantial displacement, in the order of 80 m, took place along the leading thrust in the stortorn section at this relatively late stage of development of the structures at stortorn. however, this is only a small amount of the accumulated displacement (c. 500 m) which is of the same order of magnitude as that taken up by the duplex stacking of the lower trailing-end segments of the rubjerg knude fyr and grønne rende sections. the ramping and thrusting of st01–st04 over this duplex structure explains the high elevation of the l/r-unconformity and overlying piggyback basins in the frontal part of the stortorn section. the formation of the duplex stack comprising the lower duplex segments annotated gru in fig. 113 would have taken place only after the imbricate thrusting in the grønne rende section developed (see below). the combination of displacement at the leading edge in one section and stacking of lower duplex segments in another, indicates a continuous progressive thrust-fault evolution. during the propagation of st05, the trailing end of st04 was involved in a duplex formation that resulted in fault-bend folding of the earlier formed st05 lower duplex at the stortorn lower ramp. the piggyback basin on the back of the st05 thrust sheet was deformed into a north-verging syncline due to steepening. a similar re-orientation is seen in the thrustisolated piggyback basins in st10 and st09. a marked diapirism and remobilisation of mud in the st01–st03, st05–st07 and st09 thrust sheets suggests that the diapirism was related to the intensity of ramping, especially when the ramping involved the 155 fig. 114. dynamic model of progressive deformation in the rubjerg knude fyr section illustrated in four sequential restoration cross-sections. the cross-sections demonstrate five stages in the development between the balanced cross-section and the structural cross-section (plate 2). the red lines indicate the active displacement surfaces in each deformation stage. note that the tip of the rf04 thrust sheet was displaced by normal faulting during syntectonic deposition in the rf03/rf04 piggyback basin. lower level segments and fault-bend folding of these segments. stortorn stage 6. the final ramping of lower segments from décollement level 40 m to flat level 30 m at the base of st05 terminated the translation along the lowermost 40 m décollement level. for the sections further south, the lower décollement level was situated at the 30 m level. in the frontal part of the section, continued minor compression steepened the thrust structures, and the tips of st02 and st03 were eroded and deposited in the piggyback basin of st01. stortorn stage 7. the structural complex, including the 156 st01 thrust sheet and the underlying duplex structure, became fault-bend-folded during the thrust propagation related to the progressive deformation in the rubjerg knude fyr section. stortorn section: summary data balanced length (l0): 1125 m cross-section length (l1): 570 m shortening (δl): 555 m compression: 49.3% rubjerg knude fyr section the rubjerg knude fyr section roots into the 30 m décollement level. the most striking features developed in the rubjerg knude fyr section are the large olistoliths in the piggyback basin that were derived from the collapse and gravity gliding of a projecting segment of st04. five stages in dynamic development have been distinguished, which are illustrated by four cross-sections in fig. 114. rubjerg knude fyr stage 1. during sedimentation of the first 10 m of sand of the rubjerg knude formation, the rf05, rf04 and rf03 thrust sheets were thrust up along their footwall ramps. rf04 was displaced 90 m along the upper flat level (10 m level) before the frontal part propagated up along the upper ramp. with a displacement of about 30 m, this brought the nose of the rf04 thrust sheet up into the open air, above the sedimentation level of the rubjerg knude formation. the displacement on the other two thrusts amounted to c. 20 m, implying an accumulated displacement of 70 m. rubjerg knude fyr stage 2. the exposed nose of the rf04 thrust sheet slumped down along a normal fault into the piggyback basin of rf03. at the same time, the frontal nose of rf03 was eroded away and sedimentation of the rubjerg knude formation onlapped and covered these features. at the leading edge of the section, thrusting was initiated that brought rf01 and rf02 up over what was to become the trailingend segments of the grønne rende section. rubjerg knude fyr stage 3. the frontal imbrication of rf01 and rf02 progressed during sedimentation up to about 20 m above the main l/r-unconformity level. the rf05–rf06 thrust sheet ramped up onto the intermediate flat above the trailing-end segment of rf04. the rf04 thrust sheet was displaced about 70 m up along the relatively steep footwall ramp at the trailing end of rf03. due to the fault-bend folding of rf04, the rf05–rf06 hanging-wall ramp was rotated into a vertical position. rubjerg knude fyr stage 4. when the second ‘drop’ of the frontal part of thrust sheet rf04 took place, a c. 45 m long slab of the relatively thin thrust-sheet nose slumped down along a normal fault with a vertical separation of more than 10 m. the ‘drops’ may be regarded as two break-back sequences of the rf04 thrust sheet (in the terminology used by mitra & sussman 1997; see figs 99, 100). sediment accumulation continued in the piggyback basin to a thickness of more than 30 m, including the ‘dropped’ noses of rf04. the final accumulation in the piggyback basin took place while the displacement in the rubjerg knude fyr section was concluded more than 500 m laterally to the south. the translation progressed along the 20 m flat level on top of what was to become the lower trailing-end segments of the grønne rende section. rubjerg knude fyr stage 5. the continued displacement of rf04 resulted in structural propagation of this sheet above its own piggyback basin with the ‘dropped’ thrust noses. stage 5 in the rubjerg knude fyr section is interpreted to have been contemporaneous with stage 7 in the stortorn section in which compression brought the thrust sheets into their final, steeply inclined position. rubjerg knude fyr section: summary data balanced length (l 0 ): 525 m cross-section length (l 1 ): 260 m shortening (δl): 265 m compression: 50.5% grønne rende section the impressive imbricate fan composed of 12 upright thin thrust sheets is the essential element in the grønne rende section. as a consequence of the displacement in the imbricate fan, 550 m of trailing-end lower segments were left behind to be stacked in a duplex below the frontal part of the stortorn section (fig. 113). 157 four stages have been differentiated in the development of the grønne rende section, the first three of which are illustrated in fig. 115. grønne rende stage 1. the initial thrust-fault framework was a low-angle imbrication, about 20° on each upper hanging-wall ramp, which rooted down to the upper 10 m flat level. during thrusting, the upper thrust sheets were split up into three main segments with leading thrust faults below gr02, gr06 and gr11/ gr12 which ramped down to the main level of detachment in the 20 m flat level. the initial displacement of the imbricate fan is regarded to have been 20 m on each thrust. this implies that the accumulated displacement sums up to 240 m. gr01 was not affected by thrusting in the first stage, and 240 m of its lower trailing-end segment was consequently not displaced during this stage. the sediments of the rubjerg knude formation attained a maximum thickness of 15 to 20 m during this stage, with decreased thicknesses on the back of the gr06–gr08 thrust sheets, which were elevated to the highest position. grønne rende stage 2. the imbricate thrusting progressed with a displacement of 50 m on each thrust. this implies that the hanging-wall flats were fault-bendfolded while they passed the footwall ramps, resulting in a dramatic steepening of the thrust sheets. below gr10–gr12, the gr07u and gr08u lower segments formed a duplex structure that resulted in elevation and complex ramp-propagation folding of the sheets above. the accumulated displacement implies an increase in length of the trailing-end segment of gr01 in the order of 500 m, allowing for some adjustments due to the irregular duplex deformation. sediment thicknesses in the piggyback basin in the frontal part of the section increased to 25–30 m. grønne rende stage 3. finally, the leading-edge thrust was activated and gr01 was displaced 50 m up along its footwall ramp. the gr01 thrust roots in the lower 30 m décollement level, and the displacement of the hanging-wall flat up along the footwall ramp resulted in steepening of all the early-formed thrust elements (gr02–gr13). the displacements of the individual thrust sheets range between 60 and 70 m. the thrusting resulted in the final, almost vertical, orientation of the thrust sheets. in the rear part of the section, complex deformation of the duplex below gr10–gr13 was reflected in unusual folding of the beds in the gr13 thrust sheet where folds with horizontal axial planes were formed due to gravity collapse of the piggyback basins. grønne rende stage 4. this stage concluded the thrusting of the leading hanging-wall ramp-and-flat over the footwall ramp in the trailing end of the stenstue rende section and the subsequent final rotation of the gr02– gr05 thrust sheets. in the trailing end of the section, the rf01 and rf02 sheets concluded the displacement by thrusting from the trailing-end segments of gr12 up over the footwall ramp onto the back of gr13. moreover, gr13 was rotated into an upright position whereby the horizontal axial planes became vertically orientated (plate 1). grønne rende section: summary data balanced length (l0): 1080 m cross-section length (l1): 423 m shortening (δl): 657 m compression: 60.8% comment. the lengths are measured from the footwall ramp between rf01 and gr13 to the footwall ramp between gr01 and ss06, near the thrust truncation of the l/r-unconformity. stenstue rende section two markedly different structural complexes were formed during the development of the stenstue rende section. they were mainly caused by the displacement of the same thrust sheet (ss01) when it was displaced 200 m over the upper flat on top of the piggyback basin in the sandrende section. the frontal part of ss01 above the footwall flat of the sr04 thrust sheet is one of the complexes. the other structural complex is the chaotic breccia and gravity slumping in the northern part of the section that formed as the piggyback thrust sheets were transported over a minor antiformal stack in the central lower part of the section. the progressive dynamic development in the stenstue rende section is described in terms of five stages, the first four of which are illustrated in fig. 116. stenstue rende stage 1. four minor imbrications with an accumulated displacement of 70 m initiated the 158 development in the stenstue rende section. at the leading-edge thrust, a minor connection splay separated ss02 and ss03. most of the thrusting was located at the upper 10 m flat level for a distance of about 180 m, in the northern part of which it was eventually rooted down to the lower décollement level. the displacement of the ss05 thrust sheet followed the same system, but with a smaller translation along the upper 10 m flat level. at the trailing end, ss06 was thrust up along a steep footwall ramp, and here the formation of duplex structures was probably initiated. the thickness of sediment (rubjerg knude formation) that had accumulated by this stage amounted to 10 m. stenstue rende stage 2. the leading-edge thrusting shifted to the ss01 thrust sheet, which was displaced 20 m up along the footwall ramp (the trailing end of sr04 in the sandrende section). the ss01 thrust fault extended down via an intermediate ramp to the 20 m flat level, and about 200 m from the leading footwall ramp it stepped down the lower ramp to the 30 m décollement level. at the upper hinge of the lower ramp, ss01 was folded into a fault-bend anticline, a small detachment anticline. along the foreland-dipping limb of the anticline in the ss01 thrust sheet, a normal fault was formed that displaced the tip of the ss02 thrust sheet. furthermore, the ss03 thrust sheet fig. 115. dynamic model of progressive deformation in the grønne rende section illustrated in three sequential restoration cross-sections; the final stage (4) described in the text is not illustrated. the red lines indicate the active displacement surfaces in each deformation stage. note how the shortening due to the displacement along the 20 m flat level resulted in the substantial length of the ‘left over’ lower duplex segment between the 20 and 30 m flat level. 159 became steeply inclined, and the initial imbrication of the thin ss03 thrust sheet resulted in the separation of ss03 from ss04. in the trailing end of the stenstue rende section, duplex stacking of the lower segments in ss06 resulted in elevation of the l/r-unconformity more than 10 m above the mean level. the accumulated displacement ranged up to 160 m. stenstue rende stage 3. thrusting of ss04 progressed on the upper flat over the piggyback basin of ss03 with a frontal displacement of 80 m. the hangingwall flat of ss04 ramped up along the footwall ramp of ss03 and during this translation the nose of ss05 became fault-bend-folded into a syncline with a steeply dipping southern limb. the trailing end of the ss04 thrust sheet was translated along the 10 m flat level; it was pushed from the rear by the ramping of the trailing end of the ss05 thrust sheet whereby the ss06 thrust sheet also steepened up. sediment thicknesses in the piggyback basins increased to c. 20 m, and the accumulated displacement ranged up to 240 m. stenstue rende stage 4. the dramatic major foreland thrusting of the ss01 thrust sheet, which included about 200 m displacement of the hanging-wall ramp 160 fig. 116. dynamic model of progressive deformation in the stenstue rende section illustrated in four sequential restoration crosssections; the final stage (5) described in the text is not illustrated. the cross-sections demonstrate the development stages between the initial and final positions displayed in the balanced and the structural cross-sections in plate 2. the red lines indicate the active displacement sur faces in each deformation stage. 161 over the piggyback basin of the sandrende section, occurred contemporaneously with the formation of an antiformal stack above the trailing end of ss01. the creation of the antiformal stack had already been initiated by the earlier formation of the minor detachment anticline at the ramp splitting the lower segments of ss01 (the ss01u segments). a duplex duplication of the lower ss01u segments accentuated the anticline, and finally the ss03 thrust sheet riding piggyback on ss01 was folded into an anticline with a steep foreland-dipping southern limb (fig. 90). along this limb, a normal fault developed that displaced the frontal part of the ss04 thrust sheet. a chaotic soft sedimentary fault breccia was formed during the stretching and fault separation of ss04 (fig. 91). due to an extra push from the rear, the ss05 thrust sheet was displaced a further 30 m to the south, which resulted in the formation of a huge southerly overturned slump fold above the normal fault zone (fig. 89). the accumulated displacement totals about 470 m. the displacement of the ss01 hanging-wall flat up along the steeply dipping footwall ramp constrains the sequential thrusting of the stenstue rende relative to the sandrende thrusting. thus stage 4 could not have begun before the maximum sedimentation in the piggyback basin was accomplished in the sandrende section. the initial ss01 thrusting could be regarded as a growth fault, whereby the syntectonic accumulation of sand added to the steepening of the footwall ramp. the present vertical to northerly overturned orientation of the ss01 hanging-wall flat and ramp resulted from differential thrusting and fault-bend of the ss01u lower hanging-wall ramp. note also the re-orientation of the normal fault at the tip of ss02, which due to the same deformation was bent into a horizontal position. stenstue rende stage 5. this stage corresponds to stage 6 in the sandrende section, wherein the ss01 thrust sheet riding piggyback on sr04 was displaced by normal faulting (fig. 117, stage 6). stenstue rende section: summary data balanced length (l 0 ): 760 m cross-section length (l 1 ): 285 m shortening (δl): 485 m compression: 62.5% comment. the lengths are measured from the footwall ramp between gr01 and ss06 to the footwall ramp between ss01 and sr04. if the compression was calculated from the leading-edge thrust tip of ss01 to the trailing-end footwall ramp of ss06, l1 amounts to 455 m, δl = 305 m and the calculated compression would only be 40.1%. sandrende section the dynamic development of the sandrende section was formerly interpreted as a combination of diapirism and normal faulting caused by volume exchange during thrust propagation (sadolin et al.1997). the model presented here aims at an explanation of the development purely based on a thin-skinned thrustfault model including dif ferential ramping and duplex formation. thus, the diapirism is interpreted to be an effect of ramping and fault-bend folding growth, similar to the model of mitra & sussman (1997), but also including mud-mobilisation and exaggeration of backlimb thrusting. the normal faulting occurring in the sandrende section is interpreted as the effect of differential ramping of a lower trailing-end segment that created foreland-dipping features above a hangingwall ramp propagation along an intermediate footwall flat. six stages of dynamic development have been dif ferentiated in the sandrende section (fig. 117). sandrende stage 1. after initial deposition of a 3–5 m thick succession of rubjerg knude formation sediments, the sr04 thrust sheet started thrusting about 50 m over the upper flat. the dip of the footwall ramp was relatively gentle, only c. 14°, and in the 15 m flat level the thrust fault may be traced along a minor flat segment on top of the lower trailing-end segment of sr03 (sr03u). from the minor intermediate flat, the thrust fault rooted down to the 30 m décollement level along a 20° dipping footwall ramp of sr03u. note that an upper and lower sr04 hanging-wall ramp was introduced subsequently. sandrende stage 2. translation of the lower sr04 hanging-wall ramp along the intermediate flat established the anticline in the central part of the sr04 thrust sheet. the sr03 thrust sheet started to propagate towards its foreland along the upper 10 m hanging-wall flat, and the frontal part of sr03 was displaced 50 m over the upper flat on top of the piggyback basin of sr02. the tip of the sr02 thrust sheet propagated up along a growth-fault ramp, which caused the steeply dipping 162 163 orientation of the northern boundary of the piggyback basin at the top of the sr01 thrust sheet. the accumulated displacement ranged up to about 150 m, including the initial thrusting of sr01. sandrende stage 3. during stage 3, the thickness of the sediments of the rubjerg knude formation reached 20 m in the piggyback basins in the sandrende section. in the basin at the top of the sr04 thrust sheet, the thickness varied considerably. the reason for this variation is that the top of the anticline above the sr04 lower hanging-wall ramp was subjected to erosion while deposition continued in the frontal part, south of the anticline, as well as in the basin north of the anticline. on the foreland-dipping flank of the anticline, minor sets of normal growth faults governed sedimentation (fig. 87). the tip of the sr04 thrust sheet suffered minor erosion before deposition resumed during thrust propagation. this is documented by the angular onlap relationships described by sadolin et al. (1997). sandrende stage 4. the thrusting of sr04 continued with 50 m further displacement. below the trailing end of the sr04 thrust sheet, the sr03u lower segment was picked up and displaced onto the footwall ramp of sr02. this minor duplex and ramp thrusting accentuated the sr04 hanging-wall anticline, and normal faulting on the foreland-dipping limb progressed. above the crest of the sr03u detachment anticline, a significant normal fault complex developed. here in the sr04 thrust sheet, a dense network of conjugate normal faults (fig. 85) resulted from lateral extension due to flexural slip bend over the upper hinge of the lower footwall ramp. sandrende stage 5. the thrusting of sr01 propagated up along the lower and intermediate footwall ramp of the trailing segments of the brede rende section. during this ramping, the hanging-wall flat of sr02 progressed up over the piggyback basin of sr01. the tips of the sr03 and sr04 thrust sheets thus experienced fault bending up along the footwall flat of sr02. the atypical northerly overturned tip at the top of the sr02 sheet probably formed due to accentuated reverse faulting along a former established back-thrust. in the trailing part of the section, a minor satellite splay thrust developed, which broke through the sr04 thrust sheet from the hanging-wall flat to the footwall flat below ss01. sandrende stage 6. the final development of the sandrende section was dominated by complex duplex formation and fault-bend folding of the sr01 thrust sheet below the frontal part of sr02. during the thrust propagation over the footwall ramp of the trailing-end segments of the brede rende section, a fault-bendfolded syncline was formed in sr01, which resulted in normal fault displacement of the sr01 piggyback basin and the overlying frontal part of the sr02 thrust sheet. similar normal faulting affected the ss01 thrust sheet, which had over-thrust the piggyback basin on the back of sr04. due to the intense ramping and folding of sr01 and its underlying duplex (sr01u) into an antiformal stack, mud of the lønstrup klint formation was remobilised in sr01, which intruded through the hanging-wall flat of sr02 to form the diapir in the sandrende section. sandrende section: summary data balanced length (l 0 ): 775 m cross-section length (l 1 ): 440 m shortening (δl): 335 m compression: 43.2% comment. the lengths are measured from the footwall ramp between ss01 and sr04 to the footwall ramp between sr01 and br08, at the level where the ramps cut the l/r-unconformity. the volume lost in diapirism has not been considered, and a reduced amount of compression would result by measuring l 1 from the tip of the sr01 thrust sheet to the sr04 footwall ramp. brede rende section the development of normal faults associated with foreland-dipping features of hanging-wall ramps transfacing page: fig. 117. dynamic model of progressive deformation in the sandrende section illustrated by five restoration crosssections. note that stages 2 and 3 include syntectonic sedimentation of the rubjerg knude formation, mainly related to stage 2, and the thrust-fault configuration concluding stage 3. the cross-sections demonstrate the development stages between the initial and final positions displayed in the balanced and the structural cross-sections in plate 2. the red lines indicate the active displacement surfaces in each deformation stage. 164 lated along footwall flats has already been demonstrated in the previous sections. one of the best examples of such a normal fault relationship occurs in the brede rende section. an essential element for this development was the formation of a long thrust sheet, translated laterally more than 150 m along the upper flat. this is demonstrated by the seven stages of development recognised in the brede rende section, as illustrated by the five cross-sections in fig. 118. brede rende stage 1. the first stage differentiated here is the initial sedimentation of about 3–5 m of the rubjerg knude formation. this corresponds well with the thickness of sediments deposited initially above the l/r-unconformity in the sandrende section; this unit is considered to represent pre-thrust sedimentation, i.e. the sediment record prior to piggyback basin formation. brede rende stage 2. accepting that the thinnest preserved section of the rubjerg knude formation indicates the timing of the earliest thrusting, then thrusting in the brede rende section was initiated with the displacement of the br03 thrust sheet. the frontal part of br03 was displaced about 50 m over the upper flat corresponding to the relative foreland in front of the leading edge of thrusting. the br03 thrust fault probably rooted down to the 30 m décollement level. however, translation in the upper 10 m flat level cannot be excluded, and in this case the beds disturbed by hydrodynamic brecciation might be interpreted as thrust flats. in the trailing end of the section, the br06 hanging-wall ramp was the next thrust to break through and initiate the translation along the upper flat. brede rende stage 3. the br06 thrust sheet was further displaced c. 50 m over the upper flat. the trailing end of br06 was separated by a splay thrust at the footwall ramp, along which the br07 thrust sheet propagated contemporaneously with piggyback thrusting of br08. this stage is equivalent to the frontal thrusting during stage 5 in the sandrende section. brede rende stage 4. sediment accumulation in the piggyback basins increased up to about 15 m. the marked dif ference in thickness of deposits is clearly seen by comparing the br05 thrust sheet with the br06 thrust sheet. the roof of br05 was obviously capped at an earlier stage than br06 where sediments accumulated to more than twice the thickness of that in br05. brede rende stage 5. with a displacement of about 60 m, the br05 thrust sheet propagated up along the footwall ramp of br04 and onto the upper flat on top of the br04 thrust sheet. translation of the br06 thrust sheet progressed c. 60 m along the upper flat. the thrusting rooted down to the 20 m flat level on top of the trailing-end segment of the br03 thrust sheet (br03u). the accumulated displacement amounted to 150 m, including the ramping and translation of the br07 thrust sheet along the same 20 m flat level. brede rende stage 6. after the thrusting of br05 and br06 ceased, the br04 thrust sheet was translated c. 80 m. the br04 thrust fault included three ramps: an upper gently dipping ramp from the upper flat to the 5–10 m flat level, an intermediate ramp-bend of the br03 thrust sheet due to the presence of the formerly established br02 footwall ramp, and a lower ramp from the 10 m to the 20 m flat level. the translation of the br04 lower hanging-wall ramp along the footwall flat of br03 created the foreland-dipping bend that, combined with the bend due to the br03 ramping, formed a syncline in front of the br04 ramp anticline. the normal fault created parallel to the foreland-dipping features displaced the tip of the br06 thrust sheet. the vertical offset on the normal fault amounted to c. 20 m, which also included the displacement caused by the offset in front of the br05 thrust tip. brede rende stage 7. finally, the leading-edge thrusting of the section propagated over the trailing end of the kramrende section. above the footwall ramp of br01, a minor antiformal stack was formed and subsequently an irregular duplex formation affected the br01 thrust sheet during the last stage of deformation in the brede rende section. this phase developed into diapirism that intruded towards the thrust fault between br01 and br02. facing page: fig. 118. dynamic model of progressive deformation in the brede rende section illustrated in five sequential restoration cross-sections. the cross-sections demonstrate seven stages in the development between the balanced cross-section and the structural cross-section (plate 2). the red lines indicate the active displacement surfaces in each deformation stage. note that significant normal faulting occurred in the brede rende section during stages 5 and 6 while the hanging-wall anticline in the middle part of the br04 thrust sheet was formed. 165 166 brede rende section: summary data balanced length (l0): 815 m cross-section length (l1): 440 m shortening (δl): 375 m compression: 46.0% kramrende section in the central part of the kramrende section, a major diapir developed during the progressive thrusting. the kramrende diapir was the most distally located diapir in the thin-skinned thrust-fault system indicating that a certain amount of ramp propagation from a deeper décollement level (at least 30 m flat level) was needed for macroscopic-scale diapirism. south of the kramrende section, the décollement level gradually changed to a shallower position and the intensity of ramping decreased. seven stages of dynamic development have been differentiated in the kramrende section, as illustrated in the five cross-sections in fig. 119. kramrende stage 1. the thrusting in the kramrende section was initiated with leading-edge propagation along the kr01 thrust fault, which constituted an upper footwall ramp with a dip of 10°, a minor intermediate flat at the 15 m flat level, and a c. 15° dipping lower ramp connecting the thrust fault to the 30 m décollement level. the displacement was in the order of 100 m along the upper flat, where almost no sedimentation of the rubjerg knude formation took place. kramrende stage 2. subsequent to the early stage thrusting, the lowermost 10 m of the rubjerg knude formation was deposited; the sediment thickness in the kr01 piggyback basin was probably a little less. kramrende stage 3. the kr01 thrusting progressed about 60 m over the upper footwall flat of what was to become the mb04 thrust sheet, and the trailing end of the kr01 thrust sheet was elevated to the 15 m flat level by ramp propagation over the lower footwall ramp of mb04. a small duplex segment (kr01s) under the middle part of the kr01 thrust sheet was picked up in the thrusting and displaced to the upper footwall ramp hinge, where it formed a minor angular anticline. in the syncline between the anticline and the footwall ramp of kr01, the thickness of piggyback basin sediment accumulation increased to about 15 m before the kr02 thrust sheet propagated c. 50 m up along the ramp, and the kr02 hanging-wall ramp partly capped the kr01 piggyback basin. the accumulated displacement ranged up to about 260 m. kramrende stage 4. thrusting of the kr03 thrust sheet was initiated up along the northerly dipping footwall flat of kr02. the kr03 thrust fault included an upper and a lower relatively steep (c. 23°) ramp. the top of the kr02 thrust sheet was probably exposed to erosion, and the rubjerg knude formation is thus missing in this part of the section. the piggyback sediment pile increased to a thickness of 20 m, as indicated by the sedimentary section preserved above the l/r-unconformity at the top of the kr04 thrust sheet. from the trailing end of the kr01 thrust sheet, diapirism intruded through the footwall ramp and irregular mud diapirism developed in the kr02 thrust sheet. kramrende stage 5. with a displacement of c. 30 m, kr03 thrusting propagated over the two ramps that resulted in the fault-bend folding of two anticlines separated by an intervening syncline. kramrende stage 6. the kr04 thrust sheet was thrust over the fault-bend-folds formed in stage 5, simultaneously with limited continuation of kr03 thrusting. minor irregular duplex formation started to develop into mud mobilisation at the trailing end of the kr02 and kr04 thrust sheets. kramrende stage 7. the final thrust propagation of the kr03 thrust sheet concluded with a displacement of 30 m up along the footwall flat of kr02. at the bend between the footwall flat and the footwall ramp of kr02, a remarkable set of reverse faults developed (fig. 72). the kr04 thrust sheet, carried piggyback on kr03, was also displaced by the reverse faulting, a fact that testifies to the relative timing of kr04 piggyback thrusting and kr03 ramp propagation. the reverse faults are regarded as back-limb thrusts similar to the back-thrust features mentioned in stage 5 of the sandrende section. minor back-limb reverse faults facing page: fig. 119. dynamic model of progressive deformation in the kramrende section illustrated in five sequential restoration cross-sections. the cross-sections demonstrate seven development stages, of which stage 2 represents a purely depositional phase and stage 4 only includes minor displacement. the red lines indicate the active displacement surfaces in each deformation stage. 167 168 also developed at the crest of the fault-bend-folded kr01 thrust sheet. polyphase diapirism evolved in the trailing end of the kr03 and kr04 thrust sheets, such that the primary thrust-fault framework was partially destroyed. kramrende section: summary data balanced length (l0): c. 600 m cross-section length (l1): c. 300 m shortening (δl): c. 300 m compression: c. 50% comment. the lengths are measured from approximate positions on the footwall ramps bounding the kramrende section and the data must therefore be regarded as tentative estimates. martørv bakker section the development in the martørv bakker section was dominated by the translation of a thrust sheet that was more than 600 m long and only 20–30 m thick. during nearly 400 m of displacement towards the foreland, a lower segment transformed into a duplex that ramped at a relatively late stage and created a faultfig. 120. dynamic model of progressive deformation in the martørv bakker section illustrated in three sequential restoration cross-sections. the cross-sections demonstrate four stages in the development between the balanced cross-section and the structural cross-section (plate 2). the red lines indicate the active displacement sur faces in each deformation stage. note the significant depression formed in the hanging-wall block south of the martørv bakker normal fault. in this depression, diamictites interlayered with slump-slides were deposited. 169 bend-fold anticline and syncline pair. at the upper surface of the intervening limb between the fold pair, a foreland-dipping normal fault was formed, rather similar to the structural complex formed in the brede rende section. simultaneously with the sedimentation of a diamictite, three slump-slides filled the piggyback basin developed in a syncline created at the top of the hanging-wall block of the normal fault. the sequential restoration stages are illustrated in three cross-sections in fig. 120. martørv bakker stage 1. thrusting in the martørv bakker section started with foreland thrusting of mb02, and translation of the trailing-end duplex that constituted the kr01 thrust sheet emplaced piggyback on the mb04 thrust sheet, thrust up along the footwall ramp of mb03. the more than 600 m long mb02 thrust sheet was displaced c. 105 m over the footwall flat of mb01. the mb02 thrust fault included two ramps, an upper footwall ramp of mb01 and a lower ramp between the 20 m flat level and the 30 m décollement level. the lower ramp was located below the central part of the mb02 thrust sheet, where it acted as the final step for the décollement level change to the 20 m footwall flat level. it is thought unlikely that significant sedimentation occurred in the section during this stage. 170 martørv bakker stage 2. the mb01 thrust sheet was displaced about 100 m over the foreland of the stensnæs section along the leading-edge thrust. the mb01 thrust was rooted down to the 20 m flat level, and it can be traced further on to the 30 m décollement level by passing the central lower footwall ramp of mb02. minor adjustments along the hanging-wall flat resulted in formation of small duplexes along the thrust fault. in the trailing end of the section, the mb03 thrust sheet was thrust up over the footwall ramp of mb02, whereby an antiformal stack was formed due to the folding that also involved the mb04 and kr01 thrust sheets. at the base of the mb03 thrust sheet, the lower segments formed an irregular duplex, which accentuated the antiformal stack. the accumulated displacement ranged up to 290 m. martørv bakker stage 3. the final thrusting of the martørv bakker section was concluded by nearly 100 m displacement of the mb02 thrust sheet. the frontal hanging-wall ramp-and-flat was thrust over the piggyback basin of the sn04 thrust sheet in the stensnæs section. during the thrusting, the lower segment mb02u was activated and formed a fault-bend-folded lower duplex. above the hanging-wall ramp of the trailing-end segment (mb02u3), an anticline was formed at the surface of mb02 and a subsequent syncline above the hanging-wall/footwall flat became a piggyback basin. martørv bakker stage 4. the foreland-dipping limb of the fold pair at the top of the mb02 thrust sheet developed into a southerly dipping normal fault. the c. 10 m deep piggyback basin was filled with diamictitic deposits and slump-sheets that glided down from the top of the antiformal stack. deformation in the martørv bakker section concluded with the steepening up of the leading-edge thrust structures due to ramp bending in the stensnæs section. martørv bakker section: summary data balanced length (l 0 ): 1065 m cross-section length (l 1 ): 675 m shortening (δl): 390 m compression: 36.6% comment. the lengths are measured from the tip of the leading-edge hanging-wall ramp to the upper bend of the footwall ramp of the mb04 thrust sheet. stensnæs section in the stensnæs section, a number of conspicuous flexural slip folds occur which are interpreted to have resulted from the deformation that accompanied sequential footwall ramp collapse and subsequent ramp displacement of minor duplexes. it is significant that they occur in relation to the final ramping from the lower 20 m flat level to the upper 10 m flat level. eight stages have been differentiated in the development of the stensnæs section, of which stages 2, 4 and 6–8 are illustrated by the cross-sections in fig. 121. stensnæs stage 1. in contrast to the martørv bakker section, an initial sediment thickness of 5 m of the rubjerg knude formation is thought to have covered the stensnæs section. it should be noted, however, that typical lønstrup formation facies grade upwards into typical rubjerg knude formation facies in this distal part of the rubjerg knude glaciotectonic complex; the l/r-unconformity is not clearly developed, and location of the formation boundary can be difficult. the affinities of the sediment packet referred to above are thus debatable. stensnæs stage 2. thrusting in the stensnæs section was initiated with c. 100 m displacement of the sn02 thrust sheet over the upper flat. the thrust fault ramped down to the 10 m flat level, which separated the upper and lower segments of the sn04 thrust sheet, simultaneously with stacking the sn03 thrust sheet into a northerly dipping duplex complex along the footwall ramp of sn01. stensnæs stage 3. accumulation of the rubjerg knude formation increased to a sediment thickness of 10 m. sedimentation was restricted to the piggyback basin of the sn04 thrust sheet, as well as on the foreland south of the frontal tip of the sn02 thrust sheet. stensnæs stage 4. the piggyback basin on the back of sn04 was sealed in by the overthrusting of the mb02 thrust sheet; this is equivalent to stage 3 in the martørv bakker section. stensnæs stage 5. as a trailing-end structural complex to the ulstrup section, the thrust sheets of the stensnæs section were translated together with the ul02 thrust sheet over the ulstrup footwall ramp onto the hanging-wall flat of the foreland. during ramping, the sn01 and sn03 thrust sheets were separated into small duplex segments. flexural-slip folding and polyphase 171 ul1 intermediate footwall rampsulstrup footwall ramp stensnæs 2 ulstrup footwall ramp mb1 sn4 sn4u sn4u sn1 ul2sn1 sn3 sn2 stensnæs 4 ulstrup footwall ramp mb2 mb1 sn4u sn4u sn4 sn3 sn2 sn1 sn1u ul2 ? stensnæs 6 mb1 sn4u sn4 sn3 sn2 sn 4u sn1 sn1u ul2 ul1 mb2 ? ? stensnæs 7 & 8 stensnæs ramp mb2 mb4 sn4 sn3 sn2 sn1 sn1u ul2 ul1 lønstrup klint formation active thrust fault rubjerg knude formation 100 m n s fig. 121. dynamic model of progressive deformation in the stensnæs section illustrated in four sequential restoration cross-sections. the cross-sections demonstrate five of the eight stages in the development described in the text between the balanced cross-section and the structural cross-section (plate 2). the red lines indicate the active displacement surfaces in each deformation stage. hydrodynamic brecciation resulted from the ramping (figs 53, 57, 58). the accumulated displacement ranged up to 35 m, whereas the length of the hanging-wall flat in the 10 m flat level amounted to 500 m. at the leading edge of thrusting, the ul02 thrust sheet initiated the thrusting up over a stepwise ramp. stensnæs stage 6. during this stage, about 10 m of the rubjerg knude formation was deposited in the piggyback basin at the top of the sn02 thrust sheet. the sedimentation level was probably up to 20 m above the l/r-unconformity, inferred from the elevated position of the sn02 thrust sheet. however, this is uncertain and the sediments were either never deposited or eroded away during later thrust elevation. in the northern part of the section, the sn04 thrust sheet propagated up along the footwall ramp of the earlier created sn02–sn03 duplex. this resulted in fault-bend folding of the sn04 thrust sheet and its piggyback basin as well as the overlying mb02 thrust sheet. this 172 stage correlates with stages 3–6 in the martørv bakker section. stensnæs stage 7. the sn01 thrust sheet was displaced about 50 m over its lower segment (sn01u), and together they were thrust onto the footwall ramp-and-flat of the ul02 thrust sheet. during the thrust-fault propagation of the ul02 thrust sheet over the footwall ramp of ul01, the sn01 and sn02 thrust sheets, piggyback translated on ul02, were bent into c. 30° dipping position. finally, the sn04 thrust sheet was displaced up along the footwall ramp of sn03 during dif ferential duplex formation along the sn04 hanging-wall ramp. stensnæs stage 8. the frontal parts of the sn01, sn02 and sn03 thrust sheets, as well as the anticlinal crest of the ul02 thrust sheet (formed above the upper hinge of footwall ramp of ul01), were significantly eroded, and a local piggyback basin was formed above the transition between the stensnæs and ulstrup sections. to the north of this piggyback basin, the elevated and exposed tips of the sn02–sn04 thrust sheets gravity-slumped out into the basin, where they were deposited as olistoliths, 1–5 m in size. stensnæs section: summary data balanced length (l0): 350 m cross-section length (l1): 180 m shortening (δl): 170 m compression: 48.6% 173 ulstrup section thin-skinned thrusting in the ulstrup section involved the remarkable translation of extensive, thin thrust sheets over the footwall flat of the foreland. cohesion of the thrust sheet was probably increased by ground frost in the upper part of the thrust sheet, while the hanging-wall ramp-and-flat slid on a thin zone of mobilised mud. during translation, piggyback sedimentation varied considerably. six stages have been differentiated in the development of the ulstrup section; stages 1–3 and 5 are illustrated by the cross-sections in fig. 122 (see also fig. 121). ulstrup stage 1. thrusting in the ulstrup section initiated with frontal ramping of the ul02 thrust sheet over a two-stepped footwall ramp of what was to become the ul01 thrust sheet. this ramping resulted in the formation of two, fault-propagating folded anticlines, which were separated by a shallow, broad syncline. the leading edge of the ul02 hanging-wall ramp was displaced about 25 m over the c. 5–10 m thick rubjerg knude formation deposited in the foreland (at the top of ul01). the ul02 hanging-wall flat extended along the upper 10 m flat level for about 400 m, terminating to the north at the foreland footwall ramp rooting down to the 20 m décollement level. thrust propagation up over this ramp formed a hanging-wall anticline at the trailing end of the ul02 thrust sheet. between the anticline at the trailing end and the anticline at the upper footwall ramp of ul01, a piggyback basin formed in which glaciolacustrine sediments were deposited to form the small, ephemeral ulstrup lake. fig. 122. dynamic model of progressive deformation in the ulstrup section illustrated in four sequential restoration crosssections. the cross-sections demonstrate four of the eight stages in the development described in the text between the balanced cross-section and the structural cross-section (plate 2). the red lines indicate the active displacement surfaces in each deformation stage. 174 ulstrup stage 2. thrusting along the ul02 hanging-wall ramp-and-flat progressed with an accumulated displacement of 230 m. the glaciolacustrine deposits of the ephemeral ulstrup lake participated in the ramp-propagating-folding. during translation along the upper 10 m flat level, the trailing end of the ul02 thrust sheet was probably covered by sediments, which subsequently became eroded. this event in the rubjerg knude formation corresponded to stage 6 in the stensnæs section. ulstrup stage 3. the long lateral thrusting of the ul02 thrust sheet along the upper flat resulted in 550 m of displacement, and at the lower trailing end, the hanging-wall ramp became detached to the upper footwall ramp of ul01. above this ramp, conspicuous flexuralslip folds, similar to the folds developed in the stensnæs section, were formed in the ul02 thrust sheet. ulstrup stage 4. glaciofluvial sands were deposited upon an erosional surface capping the ulstrup lake sediments (all rubjerg knude formation). depositional base level was probably equivalent to that experienced in stage 8 in the adjacent stensnæs section (see above). ulstrup stage 5. the final foreland thrusting took place as the ul01 thrust sheet was thrust over the upper ramp of the foreland and propagated about 200 m to the south. when the anticline above the ul02 hanging-wall ramp approached the footwall ramp of the foreland, where a fault-bend formed continuously during the propagation of the ul01 thrust sheet, a narrow channel was formed in which coarse-grained glaciofluvial gravel was deposited (figs 27, 122). the gravel also included redeposited frozen blocks of sand, testifying to the ground-frozen conditions of the environment (fig. 28). ulstrup stage 6. at the leading edge of the ul01 thrust fault, the deformation concluded with 150 m of displacement over the upper footwall flat of the foreland. during the translation of the ul01 thrust sheet over a minor depression in the foreland, a sandy mud volcano developed due to trapping of the high water pressure close to the leading-edge thrust. the sediment extrusion resulted in chaotic disturbances in the central part of ul01. ulstrup stage 7. the last stage of development in the ulstrup section involved sedimentation of the uppermost post-tectonic deposits of the rubjerg knude formation. the conglomerate (of stage 5) was covered by sand, and deposition in the foreland covered the leading-edge thrust at tvonnet rende (for location see plate 1). ulstrup section: summary data balanced length (l0): c. 1350 m cross-section length (l1): c. 850 m shortening (δl): c. 500 m compression: c. 37% summary of dynamic development the dynamic development of the complex is summarised in figs 123 and 124. from the scheme in fig. 123, it is clear that the rubjerg knude glaciotectonic complex developed in sequential progressive stages during syntectonic sedimentation of the rubjerg knude formation. the thin-skinned thrust-fault complex developed mainly as piggyback thrusting with proximal thrust sheets being displaced contemporaneously with activation of the distal thrust fault. during the advance of the thrust-fault complex, the position of the décollement zone shifted progressively to deeper levels. the dynamic development can be summarised in eight steps that resulted in the formation of eight characteristic thrust-fault structure types (fig. 124). fig. 123. summary scheme of the syntectonic sedimentary development in the rubjerg knude glaciotectonic complex. the deformation stages for each section, as described in the text, are indicated here by the symbol #. 175 1. long lateral translation of a thin thrust sheet took place over the foreland. the ramp was rooted in the uppermost shallow décollement level 1 at a depth of c. 10 m from the top surface. 2. ramps became rooted in décollement level 2, and the increase in ramp height, amounting to about 20 m, is regarded to be the cause of the duplex folding at ramp collapse. 3. the hanging-wall anticlines became dominant structures with hinterland-dipping piggyback thrust sheets on the back limb. as the ramps extended down into décollement level 3, the height of the ramps increased and consequently the hangingwall anticlines increased in size. 4. the antiformal stack developed, which included long-distance translated piggyback thrust sheets that were folded in a hanging-wall anticline. in relation to the antiformal stack, foreland-dipping thrust faults occur that were accompanied by normal faults. 5. the prominent imbricate fan formed above décollement level 2. the initially gently to moderately dipping imbricated thrust sheets were re-orientated into steeply dipping positions due to lateral translation of the imbricate fan along décollement level 3. 6. this step involved the subsequent deformation of the lower duplex segment not incorporated in the imbricate fan. this lower duplex segment was imbricated and the sub-segments were displaced into a duplex stack during push from behind by a progressing hanging-wall ramp, rooting in décollement level 3. 7. this step involved differential duplex stacking and imbrication of thrust-fault sheets. the piggyback basins vary in elevation due to differences in duplex stacking. furthermore, the variation in duplex stacking reflects the shift from décollement level 3 to 4 (corresponding to a shift in the décollement surface from 30 to 40 m). 8. the fault-bend-folded duplex units were formed. the formation of these duplex units was only possible because the four thrust-fault flat levels had developed, and thus the duplexes could be stacked and subsequently fault-bend-folded during maximum compression and translation along décollement level 4 (fig. 124). the thickness of sediments that accumulated contemporaneously in tectonically correlated piggyback basins decreases from north to south. thus the depocentre was situated in front of the last activated thrust section, and the depocentre gradually shifted to a more and more distal position. correlation of the syntectonic progressive development of the complex shows that sedimentation was contemporaneous with thrusting rather than there being an alternation between periods of active thrust faulting and periods of deposition. furthermore, it indicates that the ice margin was not melting back during the formation of the complex but advanced in a continuous progressive gravity-spreading process. 176 ice fault-bend-folded duplex imbricated duplex stack duplex stack imbricate fan antiformal stack hanging-wall anticline duplex folds at ramp collapse foreland translated thrust sheet distance t im e sn fig. 124. model of thrust-fault structure types formed during the progressive deformation of the thin-skinned glaciotectonic thrust-fault complex. the model outlines a progressive development in eight steps resulting in the formation of eight characteristic thrust-fault structure types, the first to develop earliest and continuously in the distal part of the complex, and the last to be formed in the most proximal part of the complex (see text for details). 177 lønstrup klint formation rubjerg knude formation stortorn formation 178 discussion the observations that form the basis for the description of the structural geology, mechanical behaviour and dynamic development of the rubjerg knude glaciotectonic complex, raise important questions with respect to understanding the framework and nature of thin-skinned thrusting related to glacial deformation; seven topics have been selected for further discussion below. the basis for understanding a structural complex is to describe the tectonic architecture and the range of structures it contains from microscopic to macroscopic scale. the discussion of thrust-fault architecture leads to evaluation of the reliability of the balanced cross-section. consideration of thrust brecciation and diapirism leads naturally to a focus on the thrustfault dynamics, and the significance of the rate of deformation. the dynamics associated with the syntectonic deposits and the formation of piggyback basins merit discussion, as does the interpretation of a proglacial contra subglacial deformational setting. the final topic deals with the geological setting of the complex, including the timing of the event that created it. thrust-fault architecture a prerequisite for understanding the thrust-fault architecture is a familiarity with the terminology (see appendix 2). the macroscopic structures encountered in thin-skinned orogenic belts are all recognisable in the glaciotectonic complex. mesoscopic structures such as folds and faults are similarly recognisable. however, small-scale structures such as joints, cleavage and fabric are more difficult to recognise (except for hydrodynamic brecciation), and this may be one of the major dif ferences between soft sedimentary deformation and hard-rock deformation. it seems likely that joints and fractures in soft sediments would be able to re-heal after deformation. thus, a large number of minor reverse faults must have formed in thrust sheets during ramp propagation (fig. 67), but appear to have disappeared again after subsequent thrust sheet propagation along the flat, as they have not been observed with the exception of the in situ positions related to ramp bend (fig. 85). there is an approximation to a right-angle relationship between the footwall ramp and the back-thrust faults, which indicates that an increase in the dip of the footwall ramp results in a decrease in the dip, in the opposite direction, of the back thrust. moreover, a steeper and higher footwall ramp also corresponds to an increase in displacement along the back-thrust fault. thus one can regard the kr01 (kramrende) back-thrust faults as structures related to initial faulting in the progressive deformation (fig. 67), and the kr04 back-thrust splay faults as a structural element related to a developed phase of progressive thrust-fault deformation (fig. 72). major back thrusting at the back of sr02 (fig. 84) represents a mature phase in the progressive thrust faulting. the apparent lack of joints and fractures reflecting ramp propagation could probably be explained as having been absorbed in the hydrodynamic brecciation process. among the structural elements analysed during the interpretation of the balanced cross-section, the duplex structures create the most interesting problems. firstly, the interpretation of the duplex imbricates in the stensnæs section provides an explanation for the complicated fold framework. secondly, the interpretation of the duplex below the frontal part of the stortorn section links the hidden duplex segments at the base of the grønne rende section with the duplex stacking below st01–st03. thirdly, the normal faults can be interpreted to have been related to the ramping of lower duplex segments. if the normal faults are regarded as foreland-dipping faults or part of a foreland-dipping duplex, the model for duplex formation suggested by contreras & sutter (1997) may be relevant for the understanding of the foreland-dipping faults. in their model for formation of forelandor hinterland-dipping duplexes, they considered two factors: u = distance of displacement along the upper flat, and s = length of duplex segment. in a regime where the ratio u/s is greater than one (u/s > 1), foreland-dipping duplexes are formed; in a regime where u/s < ½, hinterland-dipping duplexes are formed. in regimes where ½ < u/s < 1 or u/s = 1, antiformal stacks or angular antiformal stacks, respectively, are formed. this corresponds well to the interpretation presented here of the rubjerg knude cross-section, where most thrust sheets are displaced by less than their length, and consequently the main orientation of thrust sheets is hinterland dipping. according to the model of contreras & sutter (1997), foreland-dip179 ping duplexes are formed when a duplex segment is displaced along an intermediate or upper flat for a distance equal to, or more than, its length. a consequence of this is that a roofing thrust sheet will be displaced in front of the foreland-dipping upper footwall flat, where normal faulting will take place. this is interpreted to be the case for the normal faults in the martørv bakker and brede rende sections (see fig. 66). the normal fault developed in the stenstue rende section may also be regarded as an expression of the latter regime in the suggested model. a foreland-dipping feature may well reflect the foreland-dipping limb of a hanging-wall anticline formed above a laterally displaced hanging-wall ramp. however, the normal fault-displaced thrust sheet must be thrust over the duplex segment before it was thrust faulted together with its roofing thrust sheet, and then as translation continued attached to the hanging-wall flat until the displacement was concluded by the normal faulting over the tip of the duplex segment. the consideration of duplex formation naturally leads to a focus on the changes in décollement levels. when a duplex segment is formed, there would normally be an early décollement surface at a shallow level, succeeded by a shift to a deeper level connected with a new footwall ramp. the upper décollement level (the 10 m level or corresponding gently dipping ramp to the leading edge) was probably the first to be formed in the proximal part of the complex and probably also the last to form in the distal part (fig. 124). from the cross-section, it is indicated that the 10 m décollement level extended for about 1 km, but ended up with a distance of only 600 m. the 20 m décollement level was the next to take over, and during the establishment of a related ramp, footwall ramp imbrication progressed, modifying the ramp transition connecting the two décollement levels. the length of this décollement level might have been of the same scale, but only c. 400 m is preserved as a lower footwall flat. the 30 m flat level is the dominant décollement level extending from the middle of the martørv bakker section to the middle part of the stortorn section, where finally the 40 m décollement level was developed. note that the present-day position of the lowermost décollement surface is at about 45 m b.s.l. due to the regional, very gentle dip to the north of the l/r-unconformity that serves as the reference level, corresponding to the 0 m flat level. from the cross-section, it can be seen that the dips of the ramps increase from gentle (5–15°) in the zone between the upper surface flat to the 10 m flat level, to 25° dips between the 10 and 20 m levels, and reaching up to 35° between the 20 and 30 m flat levels. the ramp dips with steeper angles in the cross-section, arise from over-steepening or superimposed tilting during ramp propagation (fig. 9). thus, in the proximal part of the complex, dips between 35° and 45° are interpreted as the primary dips of ramps rooting down to the deepest décollement level at 40 m, which is incorporated in the model for the fold-imbricate duplex units. the increasing dips of ramps are interpreted to have resulted from the increase in fracture angle as a function of increase in normal stress (change of levels) and increase in shear stress (increasing force required to move thrust sheets). this is implied from the shape of the mohrenvelope in the mohr diagram (hobbs et al. 1976), and it is suggested to be a basic relationship for glaciotectonic fracture and fault deformation (pedersen 1996). balanced cross-section in the balanced cross-section, the changes in dip angles are responsible for the insertion of a number of small triangular-shaped duplex segments, which are incorporated in the geometric construction and annotated as splints (horses) (plate 2). it is not known how many splints exist in reality. a few have been recognised as structural identities (kr01 s in kramrende), but it is likely that space deficits or excesses have been absorbed in mud-mobilisation or differential small-scale anastomosing fracturing. dif ferential fracturing and thrust-fault formation with a spacing of only 1 m has been documented in the moserende section (fig. 27), indicating that the duplex segmentation does exist. hence it is probable that a much more differential translation took place than is indicated in the crosssections of the dynamic model of thrust-fault propagation (fig. 112). the reliability of the approximations in construction inherent in a balanced cross-section is founded in the area balance. the main calculation of the balance indicates that the shortening amounts to approximately 50%, with l 0 = 12 km and l 1 = 6 km. the area of the l 0 cross-section (l 0 multiplied by stratigraphic thickness) amounts to 340 000 m2, and the area of the l 1 cross-section (l 1 multiplied by measured thickness of the retrodeformed cross-section) is 382 500 m2. the dif ference amounts to 11%, which is interpreted as a consequence of the erosion of the thrust sheets in the proximal part of the complex (grønne rende section – ribjerg section). the detailed calculations of areas 180 for the area balance are summarised in table 1, and documented in plates 2a and 2b. the amount of erosion indicated from the area balance differs markedly from the 80% erosion estimate by gry (1941), and supports the argument that gry’s cylindrical thrust-fault model was incorrect. considering the amount of erosion, the question arises: why is the preservation potential so great? three factors are suggested here to answer this: (1) the steeply orientated thrust sheets were partly packed by the sand fill in the piggyback basins, (2) the thrust-fault deformation resulted in a strain hardening that consolidated the complex, and (3) as the sole of the approaching ice sheet advanced across the proximal part of the complex, the over-pressured pore water migrated from the hanging-wall ramps and flats of the thrust sheets to the hanging-wall flat of the ice sheet, facilitating the over-thrusting of the footwall block which subsequently comprised the thrust-fault complex. thrust brecciation and diapirism in order for a thrust sheet to move, a fracture must be created that can develop into a plane of thrusting. the initial fracture is formed when the failure limit is reached in a system subjected to pressure (loading and lateral compression). a recurring question, and an apparent conflict in reasoning, is why fault planes develop, leaving the rest of the thrust sheet preserved? since the sedimentary unit forming a thrust sheet is subjected to the same amount of confining pressure, it might be expected that a muddy mass of collapsed sedimentary units just as well could have been the result? it is well known that an increase in pore-water pressure results in failure and initialisation of fractures along surfaces of anisotropy, as described for orogenic systems by hubbert & rubey (1959). however, in soft sedimentary deformation with lower confining pressure and smaller shear strength, as well as smaller coefficient of internal friction, the limits of fracture formation and complete collapse are much narrower. the structures developed in the ulstrup section reflect these conditions. the anastomosing jointing and mud mobilisation at the tip of the ul01 thrust sheet reflect the stage of near collapse (fig. 49). the thick zone of hydrodynamic brecciation along the hangingwall flat reflects the same tendency towards collapse, and speculations about the influence of ground-frozen conditions on the preservation of the thin thrust sheets during translation over the foreland are relevant. ground-frozen conditions are interpreted to have affected that part of the thrust sheets elevated above the ground surface; the freezing of the sediments in the thrust sheet may result in more brittle behaviour, whereby cracks formed (figs 44, 46). however, due to the high pore pressure maintained along the hanging-wall flat, the cracks were filled with sand pumped into the cracks by the over-pressured pore water from the base of the thrust sheet. hydrodynamic brecciation is evidently related to the hanging-wall ramp-and-flat. brecciation was initiated at an episedimentary stage with the formation of ball-and-pillow structures due to sediment loading. when the loading increased by over-thrusting, the ball-and-pillow formation progressed further and hydrodynamic brecciation was concentrated at the hanging-wall flat. small-scale mud diapirism took place, with chaotic folding developing into polydiapirs (figs 54, 55, 77, 78, 86, 88). polydiapirism and mud-mobilisation are considered to have developed simultaneously and with an increasing degree of disordering and size of diapir in progressive stages of deformation. many of the mesoscopic diapir structures recognised in the rubjerg knude glaciotectonic complex can be compared with the multi-wavelength gravity structures described in the model analysis by weinberg & schmeling (1992). the formation of large-scale diapirs is suggested to have been related to thrustfault deformation of a deep-seated hanging-wall flat that propagated up to surface level along a set of relatively steep footwall ramps. during propagation up along a lower ramp to an intermediate flat, and ramping from the 20 m intermediate flat level to the 10 m flat level, polysequential hanging-wall anticlines formed, and were subsequently destroyed by mudmobilisation initiated from the deep-seated thrust zone of the hanging-wall flat. some of the soft sedimentary xenoliths floating in the mud diapirs can be viewed as relicts of anticline crests (fig. 74). staircase-like ramp propagation is indicated for the kramrende diapir and the sandrende diapir, but is not so obvious in the case of the brede rende diapir. intrusive remobilised mud is evidently related to the footwall ramp propagated hanging-wall flat of gr01, and the mud mobilisation in the thrust sheets of the stortorn and moserende sections are all easily identified with sequential ramping from the deepest décollement level. 181 thrust-fault dynamics the difference in thrust-fault development that relates to the upper flat level (10 m) is very marked when the ulstrup section is compared to the grønne rende section. thus the foreland regime in the latest stage of deformation is characterised by thin, very long sheets subjected to horizontal translation over the footwall flat of the foreland. in contrast, the grønne rende section probably formed an imbricate complex of smaller, moderately dipping thrust sheets when this section was adjacent to the foreland. there is no obvious reason for this difference, although minor differences in lithology and dif ferences in environmental conditions (frozen or unfrozen ground) could be viewed as contributing factors. however, there is an invisible condition which must be considered, namely the velocity of deformation. at the initiation of any deformation, the velocity is zero; the velocity then increases until the displacement is brought to a halt at the edge of the foreland during decreasing velocity. fast deformation results in more fractures than slow deformation. it is therefore suggested that the imbricate structures in the central part of the thrust-fault complex were initiated during the fastest advance towards the foreland and that the long-distance translation of unbroken thrust sheets relates to decreasing velocity or slow advance. in a discussion of the velocity of thrust-fault propagation, the question of rates and timing is inevitable. the youngest dating of the stortorn formation is 30 000 years b.p., while the oldest dating of the rubjerg knude formation is about 29 000 years b.p. and the oldest dating of the ribjerg formation is 26 000 years b.p. thus, a time span of 3000 years is estimated for the calculated shortening of 6 km, which indicates an average velocity of 2 m per year. the peak velocity of the deformation must evidently have been more than 2 m per year, taking into account the acceleration and deceleration. however, the velocity would also have been much higher if deformation had progressed in periodic steps rather than continuously. a step-like process would have involved periods of no movement alternating with higher velocity in the periods of advance. with respect to the rubjerg knude glaciotectonic complex, the summary of the dynamic development suggests that a continuous progressive deformation process characterised the formation of the complex (fig. 123). although the developments of the sections are described separately above, the deformational overlap from one section to the next links the sections in a continuous dynamic development. syntectonic deposition the concept of piggyback basins was originally related to large-scale regional orogenic settings (ori & friend 1984; ricci lucchi 1986). however, as applied here the term is used for the syntectonic deposits of the rubjerg knude formation that were laid down in basins structurally overlying moving thrust sheets. the initial depositional environment of the rubjerg knude formation was a relatively flat lowland, dominated by shallow lakes in an outwash plain bounded by an ice margin to the north. the plain was probably gently dipping towards the north due to isostatic loading of the ice cap. judging from the variation in thickness of the rubjerg knude formation (30 m in the proximal part to only about 10 m in the distal), the dip of the plain was not more than 2°. as the thrust belt propagated southwards, the plain became separated into smaller, more or less isolated basins characterised by steep slopes and uneven relief. the most distinctive deposits in these basins are the sedimentary breccias and slumped thrust sheets derived from the tips of up-thrust thrust sheets. three types of syntectonic slump/slide deposits can be differentiated. the first type involves deposition of coarse clasts up to metre size, which were rotated indicating transport as sedimentary clasts enveloped by sandy mud. this deposit type is regarded as being related to the distal part of the thrust-fault system, and is exemplified by the piggyback basin in the stensnæs section (fig. 56). the second type is characterised by isoclinally folded slump sheets interlayered with matrix-supported coarse clastic diamictite. this indicates that the source was very close to the depocentre, although the slump sheets were detached from their roots and were transported independently by gravity gliding into the basin. the piggyback basin in the martørv bakker section represents this deposit type (figs 23, 64). the third type comprises slump-folded sheet segments that can be traced directly, or correlated over short distances, back to the source of the thrust sheet; this type is regarded as being related to the proximal part of the system. the next step in the development would be that of thrust sheets displaced by normal faulting, but lacking depositional features such as sedimentary breccias. however, this type of dynamic development is strictly tectonic. the major 182 slump fold occurring in the stenstue rende section (fig. 89) may be regarded as a transition from a sedimentary to a tectonic regime. the deposition of recognisable thrust-sheet tips in the piggyback basins supports the concept of a continuous thrust-fault process. proglacial and subglacial deformation glaciotectonic analyses distinguish between deformation generated in proglacial and in subglacial regimes (aber 1982; croot 1988; aber et al. 1989; pedersen 1993, 1996, 2000). it has already been argued that the thinskinned thrust-fault deformation of the rubjerg knude glaciotectonic complex is an example of proglacial deformation. the key evidence for this is the presence of intimately associated syntectonic piggyback basins. these basins must have been situated in front of the ice margin, with sedimentation taking place under open water, simultaneously with thrust-fault propagation. however, the subglacial deformation is represented locally by the 1 m thick glacitectonite occurring below the glaciotectonic unconformity that truncated the thrust-fault complex. this is found in the glaciolacustrine beds at the top of the ul02 thrust sheet in the northern part of the ulstrup section. it can be argued that here the subglacial deformation penetrated down a depth of c. 5 m below the glaciotectonic unconformity. mud diapirism and hydrodynamic brecciation occurred in this setting, probably caused by loading when the ice sheet overrode the sediments. the focus of subglacial deformation is at the blå-unconformity (fig. 32). in the northernmost 300 m of the cross-section, the effects of mud mobilisation increase to a point at which primary sedimentary as well as early structural features are completely destroyed. this phase of deformation is interpreted to have taken place while the sole of the frontal part of the ice sheet was fixed to the trailing end of the thrust-fault complex. this also implies that the velocity of the thrust faulting was equal to the advance of the ice sheet. the advance of the ice-sheet load corresponds to the mechanics of gravity spreading (pedersen 1987). the increasing propagating stress resulted in increasing mud mobilisation, and subsequently the overpressure was transmitted laterally by the mud fluid towards the foreland. the mechanism might well be compared to squeezing toothpaste out of its tube. at a certain stage, the fluid pressure was released, probably due to migration of all the hydrodynamic breccias, and the mobilised mud consolidated. subsequent to consolidation, the frontal sole of the icesheet released contact with the blå-unconformity and propagated over the thrust-fault complex formed in the foreland of the ice margin. during this process, subglacial shearing affected the top of the structureless consolidated mud, and anastomosing as well as plane-parallel shear fractures were formed (fig. 32). in soft sediment structural geology, the gravityspreading model has been successfully applied to the geological setting of the mud lumps in the mississippi delta (morgan et al. 1968; pedersen 1987; aber et al. 1989). it could therefore be suggested that a gravityspreading model due to clastic progradation might be the deformation mechanism. however, there is no known delta setting at this time/place that could have provided the basis for this model, and furthermore, the sand units observed here only reach a third of the thickness of the 100 m delta-sand units in the mississippi delta setting. finally, the presence of the glaciotectonic unconformity and related glacitectonite is incompatible with a sand sediment-spreading process. although a delta setting has not been documented, it might be suggested that a slope similar to that of a megascopic delta foreset existed, and that the deformation was caused by major gravity gliding on this slope, or was simply due to uplift in the hinterland. however, this is not considered likely. the isostatic rebound documented from the elevation of the vendsyssel formation reaches 60 m a.s.l. to this must be added the uplift due to the lowering of sea level; the area in the hinterland was thus an area of subsidence rather than uplift. structurally, a gravity-gliding model would provide extensional normal fault systems in the trailing end of the thrust-fault complex (pedersen 1987). this is not compatible with the observed increase in compressional structures in the hinterland, as documented in the cross-section and indicated by the balanced cross-section; a gravity-gliding model for the complex can therefore be rejected. such larger glaciotectonic complexes are often so impressive that some geologists suggest that they were formed by orogenic activity (lykke-andersen 1992; k. binzer, personal communication 1997). disregarding the obvious glacial geological indications, there are two features that distinguish glaciotectonic complexes from basement-involved deformation: (1) the superficial detachment, and (2) the rate of translation. in the rubjerg knude glaciotectonic complex, there are no infracrustal rocks involved and the thrust sheets are not rooted down into a deep-seated hinterland source. the lowermost detachment level is 40 m be183 low the reference level, which is more or less coincident with present sea level, and there are no indications that the deformation extended below the 40 m level. the velocity of thrust-sheet motion in orogenic mountain ranges is of the order of 1 cm per year (wiltscko & dorr 1983). in glaciotectonic systems, the velocity can be up to 100 times as fast, as documented by the velocity estimate of 2 m per year for the rubjerg knude glaciotectonic complex. glacial geological conditions the rubjerg knude glaciotectonic complex is interpreted to have formed due to the advance of the norwegian ice in the late middle weichselian. the norwegian ice melted back at the beginning of late weichselian time and was succeeded by a renewed advance of the scandinavian ice sheet from central sweden. in that part of denmark east and north of the main stationary line (figs 1, 12), the direction of this advance was towards the south-west and the advance is thus referred to as the ne-ice (houmark-nielsen 1987). the eastward advance of this ice towards vendsyssel probably formed the n–s-trending hilly landscape named jyske ås, the formation of which was contemporaneous with deposition of the outwash plain represented by the ribjerg formation. when the ne-ice advance reached the rubjerg knude glaciotectonic complex, it only resulted in minor superimposed deformation. the oblique orientation of the fold axis of the megaslump in the stenstue rende section might be due to such superimposed deformation, but in general very few glaciotectonic disturbances can be related to the ne-ice advance. that overriding by the ne-ice had so little ef fect may be attributed to strain hardening due to the preceding deformation, or the smoothing out of the landscape by the former glaciotectonic unconformity, which would facilitate the second overriding of the complex. ground-frozen conditions could also have been a factor, since this would have prevented drainage from the ice sheet through the substratum, resulting in high pore-water pressures at the sole of the ice. the effect of this would have been to facilitate easy and fast propagation over the complex, although it by then formed a hill in the landscape. the norwegian ice produced a hill-and-hole pair with rubjerg knude as the hill and the depression extending from lønstrup northwards as the hole. immediately after the melting back of the ne-ice, the landscape was covered by the vendsyssel formation. a contour map of the base of the vendsyssel formation (fig. 125) thus provides a picture of the geomorphology of the young glacial landscape unaffected by the succeeding 15 000 years of erosion. in fig. 125, the hill-and-hole pair is readily identified and the general e–w morphological trends are well represented. to the east, this trend is truncated by a strong se–nw hill-and-hole geomorphology, related to the ne-ice. the trend of the thrust-fault belt of the rubjerg knude glaciotectonic complex can be followed from the coastline to the east for about 2.5–5 km. the eastern fringe of the complex has been eroded down to sea level, probably by the ne-ice, and subsequently concealed by the vendsyssel formation. 184 185 facing page: fig. 125. contour map of the pre-vendsyssel formation landscape; for location, see fig. 13. note the depression north of lønstrup which represents the hole in the hill-and-hole pair morphology of a glaciotectonic complex; the corresponding hill is represented by the high at rubjerg knude. the map is based on data from the geus well database and from plate 1. conclusions structural analysis of the rubjerg knude glaciotectonic complex, based on detailed photogrammetric measurements and field investigations, provides a geological cross-section through a low-friction thrustfault system. interpretation of the entire thrust-fault architecture included unexposed parts of the complex, and is based on the construction of a balanced crosssection. a model for the dynamic development demonstrates that deformation progressed continuously and involved formation of duplexes and mud diapirs. although the thrust-fault structures were formed in a proglacial regime related to the advance of the norwegian ice (30 000 – 26 000 b.p.), the structures can be viewed as representing an almost complete model of thin-skinned thrust-fault systems. for descriptive purposes, the complex is subdivided into 13 sections, which demonstrate the structural development from a proximal to a distal position in the thrust-fault system. investigation of these sections provided the following main conclusions. 1. the structural elements in the rubjerg knude glaciotectonic complex comprise ramps and flats related to hanging-wall and footwall positions, respectively. hanging-wall anticlines and footwall synclines were formed due to thrust-fault propagation. back-thrust faults were formed during upper ramp-hinge propagation, and an irregular fold framework developed in relation to sequential duplex imbricate formation during footwall ramp collapse. foreland-dipping normal faults were formed in relation to translation of duplex segments. 2. from the balanced cross-section, the shortening during thrust-fault deformation is calculated to have been c. 50%. about 11% of the initial stratigraphic unit subjected to thrust faulting is estimated to have been lost due to erosion. the décollement zone was at its deepest position (40 m) in the proximal sections, becoming shallower towards the foreland. stacking of duplex segments is correlated with space problems created in the subsurface due to initial displacements at the upper levels. stacking of duplex segments correlates well with the elevation of the reference level in the system. 3. hydrodynamic brecciation was dominantly related to the hanging-wall ramps and flats. polydiapirism and mud mobilisation characterise the thrust zones. mud mobilisation resulted in the formation of larger mud diapirs, and preferentially evolved during hanging-wall propagation from the décollement level up above sets of intermediate and upper footwall ramps. 4. syntectonic deposition took place in piggyback basins overlying the thrust sheets. thrust sheets exposed to erosion provided sediment to the basins, and in some cases major lumps derived from the tips of thrust sheets slumped and slid as megablocks into the piggyback basins. 5. the thrust-fault deformation was caused by gravity spreading at the front of an advancing ice sheet. over-pressured mud formed an important part of the stress transfer. the average velocity of the thrust-fault displacement is estimated to have been 2 m per year. a 40 m thick succession of flat-lying sediments, extending for 12 km, was compressed into a thrust-sheet complex that was 6 km in length and up to 80 m thick. acknowledgements the geological survey of denmark and greenland is thanked for supporting this project during the last 10 years. initial investigations were carried out while the author held a senior stipend at the geological institute, university of copenhagen. the carlsberg foundation supported the project with a one year research grant, which is gratefully acknowledged; the danish research agency is thanked for financial support for printing this bulletin. keld dueholm and the institute of survey and photogrammetry are thanked for their co-operation and willingness to provide time and fa186 cilities at the photogrammetric instrument at the technical university of denmark. frants von platen-hallermund is thanked for assistance with the arc-info transformation and arc-view editing, which provided the graphic display of plates 1 and 2. alice rosenstand and benny m. schark helped in drafting the figures, and the staff of 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: a surface along which an overlying block is displaced relative to an underlying block. relative to bedding, two different elements are distinguished in a thrust fault: the ramp and the flat. ramp: a thrust-fault ramp cuts up-section in the direction of slip and dips towards the hinterland. the angle between bedding and the ramp is in general between 20° and 30° and will not exceed 45° due to general rules of initial fracturing. a ramp is linked to a lower flat at the lower ramp hinge and to an upper flat at the upper ramp hinge. a ramp may become foreland-dipping in special cases, mainly related to transport along an upper flat. flat: a thrust-fault flat is a bedding-parallel slip surface along which lateral displacement takes place. the lowermost thrust-fault zone in deformation complexes is in general referred to as the décollement surface, décollement zone or décollement level. in the present description the ‘décollement level’ is the term used for the thrust fault between a thrust sheet and an undisplaced footwall block below a footwall flat. as flats develop at dif ferent levels, the flats above the décollement level are referred to as intermediate flats and the upper flat (identical with the roof thrust fault). the thrust-fault flats are referred to by their depth from the upper reference zero-level indicated from the balanced section. thus the 20 m flat level is the horizontal thrust fault situated 20 m below the top reference level and 10 or 20 m above the décollement level. thrust sheet: a thrust sheet is the block displaced over a thrust fault. in this study, the thrust sheets are annotated according to the section in which they occur with two capital letters, referring to the named section, and a number referring to its position from leading edge to trailing end of the section. thus, kr01 is the thrust sheet nearest to the foreland in the kramrende section. a thrust fault is referred to according to the thrust sheet it displaces. a thrust sheet is synonymous with the hanging-wall block. hanging-wall block : the rock mass displaced over a thrust fault is a hanging-wall block. at the base, a hanging-wall flat and a hanging-wall ramp bound the hanging-wall block. at the roof, the hanging-wall block is capped by a top surface or a roof thrust fault. the roof thrust fault may constitute a footwall flat as well as a footwall ramp. footwall block: the rock below a thrust fault is a footwall block. the footwall block is bounded by a footwall ramp, and the top of the footwall block constitutes a top surface and/or a footwall flat. hanging-wall ramp: the segment of a ramp that bounds the hanging-wall block is a hanging-wall ramp. at the incipient displacement along a ramp, the hanging-wall ramp is thrust along a footwall ramp. when the hanging-wall ramp passes the upper ramp hinge, the hanging-wall ramp is thrust along a footwall flat. a hanging-wall anticline is always formed above a hanging-wall ramp. hanging-wall flat: the bedding-parallel thrust-fault boundary below the hanging-wall block is a hangingwall flat. when a hanging-wall flat is thrust up along a footwall ramp, the hanging-wall flat is re-orientated and becomes inclined towards the hinterland. when the hanging-wall flat is thrust along an upper footwall flat, the thrust fault again becomes bedding parallel. footwall ramp: the inclined thrust-fault boundary of a footwall block is a footwall ramp. the footwall ramp is either the ramp boundary to the undisplaced foreland or it forms the trailing ramp boundary of a thrust sheet. in this study, the trailing footwall ramp is referred to using the annotation of the thrust sheet/ footwall block that underlies it. thus the kr02 hanging-wall ramp is displaced up along the kr01 footwall ramp. footwall flat: a footwall flat is always the top of a footwall block. a footwall flat is more or less horizontal unless it is re-orientated during the displacement of a thrust sheet up along a ramp. hanging-wall anticline: when a hanging-wall block is thrust over an upper ramp hinge, the hanging-wall block is folded into an anticline with a foreland-dipping forelimb and a hinterland-dipping backlimb. this fold may also be termed a ramp anticline. during the progress of thrusting along the upper limb, the hanging-wall anticline develops into a flat-topped anticline. the flat-topped anticline may alternatively be regarded as a flat-lying thrust sheet with a foreland-dipping forelimb or a frontal thrust-sheet nose. however, it is important to note that above a hanging-wall ramp thrust along a footwall flat, a foreland-dipping surface is formed. 191 footwall syncline: when a thrust fault propagates up along a ramp, an anticline–syncline pair is formed above, and in front of, the tip of the thrust fault, identical to the formation of a fault-propagation fold. when the thrust fault finally breaks through the folded layers, the fold pair is separated into a hanging-wall anticline and a footwall syncline. a footwall syncline therefore represents the gentle deformation below the footwall ramp; this deformation does not add significantly to the displacement along the thrust fault. the footwall syncline may also be regarded as a drag fold. the case where this is the only correct interpretation is along a growth fault. here the sediments deposited syntectonically up against a hanging-wall ramp are successively bent into an overturned syncline. the footwall syncline is identical to a trailing syncline. duplex: a duplex is one or more thrust-sheet segments entirely bounded by thrust faults and thus overlain by a thrust sheet. a thrust sheet bounded by thrust faults is called a horse, originally regarded as a minor rootless thrust-sheet segment. some of the lower thrustsheet segments described in this study are identical to horses, although the more neutral term ‘segment’ is adopted here. the formation of a duplex is related to the ‘footwall ramp collapse’ (boyer & elliott 1982), whereby progressive failure during thrust-fault propagation creates successively younger thrust faults below older ones. a parcel of thrust-sheet segments may be stacked to form a duplex complex. imbricate fan : a branching thrust-fault complex in which the individual thrust faults reach the surfaces or top level is called an imbricate fan. an imbricate fan is termed a duplex if the upper boundary is a roof thrust. antiformal stack: when a duplex is fault-bend-folded over a footwall ramp, an antiformal structure similar to a hanging-wall anticline is formed. due to the complex stratigraphic relationship within such a structure, it is referred to as an antiformal stack. piggyback thrusting: when an older thrust sheet rests on the back of a younger thrust sheet and is transported due to the displacement along the thrust faults bounding the younger thrust sheets, it is called piggyback thrusting. piggyback basin: just as piggyback thrusting refers to transport of a thrust sheet, the term is also applied to a basin that accumulates sediments during translation on the back of an active thrust sheet: the piggyback basin (ori & friend 1984; ricci lucchi 1986). in this study, the term is mainly used in the description of an area of sedimentation between two thrust sheets. in general, the piggyback basin is deposited between a fault-bend thrust-sheet tip in the distal part of a thrust structure and bounded by a hanging-wall ramp at the proximal boundary of the basin. the term piggyback basin can only be applied to successions identified as having been deposited syntectonically. 192 appendix 2 specification of photogrammetric work the construction of the rubjerg knude cross-section (plate 1) is based on a multi-model photogrammetric investigation of the cliff section, with the application of the method described by dueholm (1992). a series of oblique photographs were taken from a cessna fixed-wing aircraft in june 1993. the camera used for the photography was a minolta xg2, which had been tested and calibrated for its optical specifications at the laboratory of photogrammetry at the danish technical university. the films used were standard 24 × 36 mm colour diapositive. the photographs were taken with 66% overlap from a distance of 200–300 m with an inclination angle of c. 35°. from the series of photographs, 70 samples were selected for setting up three sets of templates, which included 67 stereoscopic models. in the laboratory, the orientation of the stereo-models was carried out based on ground control points adapted from two sets of vertical aerial photographs at a scale of 1:25 000, namely d9202 g 1365–66 and kms 9203 a509–10 taken in may 1992. the strike of the section line is n15°e from rubjerg knude and southwards. north of rubjerg knude, the strike is n24°e, which is nearly parallel to the direction of the coastline along the beach. fortunately, this is also a reasonable approximation of being perpendicular to the main concentration of structural strikes (bedding, thrust faults and fold axes; fig. 10). a minor adjustment of the northern and southern section lines was subsequently implemented to make the cross-section fit to the general plane of orthographic projection with a projection axis striking 107°. the stereoscopic instrument used for the investigation was a kern dsr 15 analytical plotter with a dec vms operating system and the special attached geoprogram developed by dueholm (1992). five different labels were used for the features outlined by the floating mark: line type 1 includes bedding traces, line type 2 includes the main unconformities, line type 3 was used for the contacts between geological units (members and formations), line type 4 outlines thrust faults, and finally line type 5 was used for topographic features (dunes, scree cones, strandplain, rockfalls etc.). digitalisation of the geological structures in the stereo-models was administrated in data files, each covering a plot-area. the plot-areas covered 500 m of the rubjerg knude cross-section, and 13 plot-areas were used for the analogue plotting of data digitised in the stereo-model. the digital data were stored for the later construction of the cross-section and the transformation for other programs applied for the management of the cross-section display. the orientation of models and setting up the system for the cross-section investigation took about one week, and the photo-geological compilation work was made over a period of three months in the autumn of 1993. the average progress was two models per day. the benefit of the multi-model analytical stereo-plotter is that features can be traced continuously from one model to the adjacent models. thus one is not restricted to working model by model, but the compilation can be extended over several models using the same set of templates. by january 1994, the cross-section could be plotted out in a normal vertical projection profile plan from the stored digital data with the application of the program facilities prepared by dueholm (1992). the scale of the rubjerg knude crosssection in the draft versions is 1:500, and the accuracy of the plotted data is estimated to be better than 25 cm. in 1995–1996, the cross-section details observed in the photo-geological models were checked in the field, and in 1997 the templates were set up again for correction, adjusting and compilation of details in the cross-section. rosa-2006:rosa-12.5 10/13-kopi a gold prospect on central storø in the nuuk region of southern west greenland is hosted by a sequence of intensely deformed, amphibolite facies supracrustal rocks of late mesoto neoarchaean age. the prospect is at present being explored by the greenlandic mining company nunaminerals a/s. amphibolites likely to be derived from basaltic volcanic rocks dominate, and ultrabasic to intermediate rocks are also interpreted to be derived from volcanic rocks. the sequence also contains metasedimentary rocks including quartzites and cordierite-, sillimanite-, garnetand biotite-bearing aluminous gneisses. the metasediments contain detrital zircon from different sources indicating a maximum age of the mineralisation of c. 2.8 ga. the original deposition of the various rock types is believed to have taken place in a back-arc setting. gold is mainly hosted in garnetand biotite-rich zones in amphibolites often associated with quartz veins. gold has been found within garnets indicating that the mineralisation is pre-metamorphic, which points to a minimum age of the mineralisation of c. 2.6 ga. the geochemistry of the goldbearing zones indicates that the initial gold mineralisation is tied to fluid-induced sericitisation of a basic volcanic protolith. the hosting rocks and the mineralisation are affected by several generations of folding. regional geology the nuuk region is renowned for the presence of rocks representing at least three major episodes of crustal accretion during the first billion years of the earth’s history. each episode is characterised by early formation of supracrustal rocks followed by intrusion of tonalites now occurring as grey hornblende-biotite gneisses. the rock packages have subsequently been deformed and metamorphosed, resulting in often very complex outcrop patterns. hollis et al. (2006) outlined the following sequence of events. the eoarchaean isua supra crustal belt (3.87–3.71 ga) is cut by 3.85 to 3.65 ga tonalites deformed and metamorphosed at c. 3.65 ga. the next sequence of supracrustal rocks was formed at c. 3.1–3.07 ga, intruded by tonalites at c. 3.07–3.05 ga, and deformed and metamorphosed at c. 2.98 ga. the supracrustals linked to the third episode were formed at c. 2.85 to 2.8 ga, intruded by tonalites at c. 2.83 to 2.75 ga, and deformed and metamorphosed at 2.7 to 2.6 ga. the last major event in the region was the intrusion of the qorqut granite at c. 2.53 ga. the rocks formed during these episodes crop out in a complex pattern of blocks or terranes often only a few tens of kilometres wide. attempts have been made to map and name these terranes (e.g. nutman et al. 2004; friend & nutman 2005), but as the rock types belonging to the different epi sodes are very similar, it has proved to be very difficult to distinguish and map the terranes in the field (hollis et al. 2006). gold-hosting supracrustal rocks on storø, southern west greenland: lithologies and geological environment christian knudsen, jeroen a.m. van gool, claus østergaard, julie a. hollis, matilde rink-jørgensen, mac persson and kristoffer szilas © geus, 2007. geological survey of denmark and greenland bulletin 13, 41–44. available at: www.geus.dk/publications/bull 41 fig. 1. map of supracrustal units in the central part of storø. the supra crustal rocks are cut by numerous pegmatites not shown on this map. the gold-hosting supracrustal sequence overlies the storø shear zone (figs 1, 2), an oblique ductile thrust which juxtaposes the supracrustal sequence against the c. 3050 ma tona litic nûk gneisses to the west. the gneisses to the south-east are thought to be palaeoarchaean tonalites. the thrusting along the storø shear zone occurred syn-to-post the latest folding event and is broadly coeval with a phase of pervasive intrusion by pegmatites at 2630 ma. lithologies the supracrustal rocks on storø comprise a range of lithologies including quartzites, quartzo-feldspathic gneisses, aluminous schists and gneisses, amphibolites and ultramafic rocks. the distribution and relative abundance of rocks in the central part of the island is shown in fig. 1. the rocks are metamorphosed to amphibolite facies, and the dominant litho l ogy is mafic amphibolite. the amphibolites (mafic, foliated, amphibole-dominated rocks) can be divided into a number of types. they are all highly strained while compositional layering varies from subtle variations on dmto m-scale in dark amphibolite, to more distinct cm-scale compositional layering in more leucocratic amphibolites. the most common type is homogeneous amphibolite, a mediumto coarse-grained, black to dark grey rock mainly composed of hornblende and plagioclase. the proportion of plagioclase to hornblende varies, and minor amounts of quartz, biotite, sphene, apatite and ilmenite occur. locally, near some large anorthosite bodies at the struc tural base of the sequence, the hornblende or the plagioclase in the amphibolite occurs as cm-sized aggregates, suggesting that the amphibolite here may represent metamorphosed gabbroic rock. varieties of the homogeneous amphibolites include garnet amphibolites and diopside amphibolites. several types and generations of quartz veins have been observed in the amphibolites. these include (1) 1–2 cm thick sheeted quartz veins, with sharp contacts to the host rocks, concordant and generally isoclinally folded; (2) thin, spidery quartz veins anastomosing through the rocks; (3) centimetre to decimetre thick agmatitic quartzo-feldspathic veins that anastomose within 0.5–2 m wide zones; and (4) generally concordant quartzo-feldspathic veins, commonly with minor garnet. a suite of foliated felsic and aluminous gneisses is interleaved with the amphibolites on storø (fig. 2). near the structural base of the supracrustal sequence, grey compositionally layered felsic biotite-hornblende gneisses of intermediate composition are often seen. the layering is expressed by variable proportions of hornblende, biotite, quartz and plagioclase. sphene, apatite and ilmenite are common accessory minerals while garnet is uncommon. the aluminous gneisses contain high and variable amounts of garnet, biotite, sillimanite and cordierite, and minor amounts of staurolite and tourmaline. these rocks are characterised by strong foliation combined with compositional layering (cmto dm-scale) to lamination (mm-scale; fig. 3a). amongst these rocks, a garnet-biotite gneiss, a sillimanite-garnet-biotite gneiss and a cordierite-garnet-biotite gneiss can be mapped as lithological units. these gneisses are characterised by a rusty brownish weathering colour (fig. 2) mainly due to the high content of biotite. within the package of aluminous gneisses, a c. 1 m thick characteristic and mappable unit with a very high content of magnetite and garnet occurs (fig. 1). a magnetite-free quartz-garnetite, consisting of up to 90% garnet is interlayered with the magnetite-bearing rocks. where the degree of exposure is high and using the garnet-magnetite rock as a marker horizon, a tight to isoclinal fold pattern with repetition of the tectono-stratigraphic units can be identified in the biotite-garnet gneiss. within these aluminous gneisses, a unit dominated by quartzite occurs. the composition varies from pure quartzite (>90% quartz) to a rock composed predominantly of quartz with feldspar, sillimanite and light green muscovite, and locally with minor fuchsite, biotite or garnet. at the other end of the spectrum, ultramafic lenses (boudins) form a characteristic component in the supracrustal sequence. the size of the lenses varies from round pods tens of metres in diameter to elongate lenses hundreds of metres thick and up to 1 km long (fig. 1). the modal composition of the ultramafic rocks varies substantially. in the cores of the major bodies the rock is mainly composed of olivine, pyroxene (both orthopyroxene and clinopyroxene) 42 fig. 2. view of the storø supracrustal units looking north-east from qingaq. to the left, the south-west-dipping storø shear zone is visible. the rusty weathering rocks are supracrustals mainly thought to be of metasedimentary origin. the amphibolites are dark grey and the anorthosite body light grey. highest summit at left is 1425 m. and amphibole (hornblende and tremolite). the margins of these bodies consist commonly of tremolite-phlogopite schist with variable chlorite and actinolite. this rock is locally affected by prograde breakdown of chlorite producing a conspicuous texture where large aggregates of parallel platy olivine crystals are intergrown with large chlorite crystals at high angles to the foliation (fig. 3b). the ultramafic rocks are locally altered to serpentinite. gold mineralisation gold occurrences are located in the amphibolite-dominated parts of the supracrustal sequence. they are often found in zones where the amphibolite is enriched in biotite and garnet and often associated with pyrrhotite, arsenopyrite and loellingite (feas2). gold is found both in sericitised plagioclase and within garnet (juul-pedersen et al. in press), indicating that the gold mineralisation predates the metamorphic event at c. 2700–2630 ma (hollis 2005). the gold is also often associated with quartz veins, and in these visible gold is locally found. geochemistry the amphibolites have a composition equivalent to tholeitic basalts and are interpreted as metamorphic equivalents of basalts. based on their geochemistry, polat (2005) stated that the amphibolites have a subduction zone geochemical signature. the common occurrence of calc-silicates is taken as a sign of hydrothermal alteration in a sea-floor environment (polat et al. 2007). the grey biotite hornblende gneiss of intermediate composition could be the metamorphic equivalent of an intermediate volcanic rock. the ultramafic rocks have mgo in the range of 20 to 40 wt%, high cr and ni (1500 to 4000 ppm) and cao/al2o3 ratios around 1, suggesting a komatiitic protolith. the garnet-biotite gneiss, cordierite-garnet-biotite gneiss and sillimanite-garnet biotite gneiss are generally characterised by high al contents consistent with a sediment protolith, which is supported by the observation of detrital zircon grains in these rocks (hollis et al. 2006). the gneisses of supposed sedimentary origin have distinctly different zr/tio2 ratios compared to the amphibolites (fig. 4). there are, however, some extremely aluminous rocks (consisting primarily of garnet and sillimanite together with quartz, feldspar and biotite) in contact with the amphibolite on qingaq, and these rocks have zr/tio2 ratios similar to the amphibolites (fig. 4). as neither zr nor ti is likely to be mobile in a non-alkaline environment they have probably been derived from basalt by sericitisation caused by hydrothermal alteration (cf. garde et al. 2007 – this volume). this process can lead to volume loss due to loss of elements such as ca, si, fe and mg, and accordingly to increased concentrations of al, ti and other immobile elements together with k (stable in the sericite). this rock type contains gold likely to have been introduced during the intense (premetamorphic) hydrothermal alteration. the more common goldbearing zones within the amphibolites are characterised by elevated contents of biotite and garnet. these rocks have increased concentrations of al and k likely to have been caused by a similar process. ages and structural relations detrital zircon from a garnet-biotite-sillimanite-cordierite gneiss and from a quartzite has been analysed (hollis 2005; rink-jørgensen 2006) using laser ablation techniques (cf. frei et al. 2006). the results indicate that there are two distinct sources for these rocks (fig. 5). the garnet-biotite-sillimanite-cordierite gneiss contains detrital zircon distributed 43 fig. 3. appearance of selected lithologies. a: thinly laminated, rusty, garnet-sillimanite-biotite gneiss with tight to isoclinal folds. b: ultramafic rock with a conspicuous texture consisting of sets of parallel oriented, up to 10 cm large, platy, olivine crystals together with large chlorite crystals. in two age populations around 2880 and 2830 ma. we suggest that the rock had an immature sedimentary precursor, possibly derived from erosion of a relatively young volcanic arc formed in two episodes at 2880 and 2830 ma. the detrital zircon in the quartzite, which is likely to represent a mature sediment, forms populations with main peaks around 2960 and 3010 ma and a minor peak at 3180 ma. this rock must have had a distinctly different source and probably represents the erosional products of an older continent (e.g. the nûk gneisses). this difference in sources is consistent with a model of formation of the supracrustal rocks in a back-arc environment. the immature sediments are likely to have been derived from the relatively young rocks in the arc, whereas the mature quartzites are likely to originate from the older rocks located on the continent side of the back-arc basin. this is supported both by the observation that the amphibolites formed in an arc environment and by the suite of rocks present. this environment is very similar to that envisaged for canadian neoarchaean supracrustal belts (sandeman et al. 2006). references frei, d., hollis, j.a., gerdes, a., harlov, d., karlsson, c., vasquez, p., franz, g. & knudsen, c. 2006: advanced in situ geochronological and trace element microanalysis by laser ablation techniques. geological survey of denmark and greenland bulletin 10, 25–28. friend, c.r.l. & nutman, a.p. 2005: new pieces to the archaean terrane jigsaw puzzle in the nuuk region, southern west greenland: steps in transforming a simple insight into a complex regional tectonothermal model. journal of the geological society (london) 162, 147–162. garde, a.a., stendal, h. & stensgaard, b.m. 2007: pre-metamorphic hydrothermal alteration with gold in a mid-archaean island arc, godthåbsfjord, west greenland. geological survey of denmark and greenland bulletin 13, 37–40. hollis, j.a. (ed.) 2005: greenstone belts in the central godthåbsfjord region, southern west greenland: geochemistry, geochronology and petrography arising from 2004 fieldwork, and digital map data. dan marks og grønlands geologiske undersøgelse rapport 2005/42, 215 pp. hollis, j.a., frei, d., van gool, j.a.m., garde, a.a. & persson, m. 2006: using zircon geochronology to resolve the archaean geology of southern west greenland. geological survey of denmark and greenland bulletin 10, 49–52. juul-pedersen, a., frei, r., appel, p.w.u., persson, m. & konnerupmadsen, j. 2007: a shear zone related greenstone belt hosted gold mineralisation in the archean of west greenland. a petrographic and combined pb-pb and rb-sr geochronological study. ore geology reviews 32, 20–36. nutman, a.p., friend, c.r.l., barker, s.l.l. & mcgregor, v.r. 2004: inventory and assessment of palaeoarchaean gneiss terranes and detrital zircons in southern west greenland. precambrian research 135, 281–314. polat, a. 2005: geochemical and petrographic characteristics of the ivisartoq and storø greenstone belts, southern west greenland: progress report. danmarks og grønlands geologiske undersøgelse rapport 2005/42, 80–112. polat, a., appel, p.w.u., frei, r., pan, y., dilek, y., ordónez-calderon, j.c., fryer, b., hollis, j.a. & raith, j.g. 2007: field and geochemical characteristics of mesoarchean (~3075 ma) ivisartoq greenstone belt, southern west greenland: evidence for seafloor hydrothermal alteration in supra-subduction oceanic crust. gondwana research 11, 69–91. rink-jørgensen, m. 2006: in situ u-pb datering af zirkoner i teori og praksis til bestemmelse af alder og populationer af detritale zirkoner, med anvendelse på en prøve fra nuuk regionen i det sydlige vestgrønland. unpublished b.sc. project, københavns universitet, danmark. sandeman, h.a., hanmer, s., tella, s., armitage, a.a., davis, w.j. & ryan, j.j. 2006: petrogenesis of neoarchaean volcanic rocks of the macquoid supracrustal belt: a back-arc setting for the northwestern hearne subdomain, western churchill province, canada. precambrian research 144, 140–165. 44 authors’ addresses c.k. & j.a.m.v.g., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: ckn@geus.dk c.ø., nunaminerals, vandsøvej 5, p.o. box 790, dk-3900 nuuk, greenland. j.a.h., northern territories geological survey, p.o. box 3000, darwin nt 0801, australia. m.r.-j., m.p. & k.s., institute of geography and geology, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. fig. 4. plot of zr ppm vs. tio2 wt% in different rock samples from qingaq on storø. fig. 5. detrital zircon populations from two rock samples from storø. sample 481465 is a garnet-biotite-sillimanite-cordierite gneiss and 481283 is a quartzite. geological survey of denmark and greenland. bulletin 10, 53-56 53 in 2006 an important milestone will be reached in the study of the three-dimensional structure and architecture of the nuussuaq basin in west greenland. the fifth and last of a series of detailed geological profiles through the sedimentary and volcanic rocks of the nuussuaq basin on disko and nuussuaq will be published (fig. 1). at the same time, the last geological map at scale 1:100 000 of the basin area will be completed. these studies have been carried out over more than two decades by a group of scientists within geocenter copenhagen and the technical university of denmark. the five geological profiles are at scale 1:20 000 and have been published as coloured plates in the same format as the geological maps (pedersen et al. 1993, 2002a, 2003, 2005, 2006). in total, the profiles cover about 500 km of cross-sections through a classic sedimentary-volcanic basin and its crystalline basement. this is one of the best exposed basins of its kind on earth, and it serves as a reference area for studies of similar basins on the continental shelves of greenland, north-western europe and elsewhere. the nuussuaq basin during the late mesozoic, a series of linked sedimentary basins were established along the west greenland continental margin. the nuussuaq basin is the only one of these basins that extends into the onshore areas where late cretaceous to paleocene sediments overlain by volcanic rocks can be studied in detail. from the late albian, sandstones and shales were deposited in a fluvial to deltaic environment in eastern disko and central nuussuaq. towards west and north the delta fanned into deeper water. several tectonic phases affected the basin and gave rise to an uplifted rift margin of precambrian gneiss in eastern nuussuaq and a gneiss ridge in central disko (chalmers et al. 1999). in the paleocene around 62 ma (storey et al. 1998), volcanism started on the sea floor in the western part of the area. the volcanic pile rapidly became emergent and the lavas prograded eastwards into the marine embayment and gradually filled it with thick foreset-bedded hyaloclastites capped by subaerial lava flows. thus, most of the early volcanic units exist in both subaerial and subaqueous facies. eventually the marine basin, and also some slightly younger lake basins, was obliterated. the volcanic rocks interfingered with fluvial sandstones and lake sediments in southern nuussuaq and eastern disko, whereas in eastern nuussuaq they lapped onto the precambrian highlands. the volcanic rocks form two formations: the older vaigat formation of thin pahoehoe flows of magnesia-rich picrites and their equivalent hyaloclastites, and the overlying maligât formation of thick, extensive lava flows of basalt and their five slices through the nuussuaq basin,west greenland asger ken pedersen, lotte melchior larsen, gunver krarup pedersen and keld s. dueholm © geus, 2006. geological survey of denmark and greenland bulletin 10, 53–56. available at: www.geus.dk/publications/bull 54°w 52°w 69°30'n 70°30'n 70°n nuussuaq vaigat mali ga at disko saqqaq pingu ikorfat itilli fig.2 fig.4 fig.3 25 km maligât formation vaigat formation cretaceous–palaeogene sediments basement gneiss fault ice cap uummannaq aaffarsuaq qeqertarsuaq ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■■ 3 3 4 4 1 1 2 2 5 5 5 fig. 1. index map, with red lines showing the location of the five geological profiles through disko and nuussuaq. profile 1: pedersen et al. (1993); profile 2: pedersen et al. (2002a); profile 3: pedersen et al. (2003); profile 4: pedersen et al. (2005); profile 5: pedersen et al. (2006). equivalent hyaloclastites. a detailed stratigraphy of the volcanic succession has been established based on lithology and geochemical analyses; several distinct units of crustally contaminated rocks form particularly good marker horizons because they are interbedded with the normal lithologies. oil seepages are found in the volcanic rocks in western nuussuaq and northern disko, and the nuussuaq basin is a key area in the exploration for oil both onshore and offshore west greenland. exploration for nickel and platinum has been focused on some of the contaminated volcanic units in which these elements are missing from the surface deposits and may have formed accumulations at depth in magma chambers. methods many mountain walls in the nuussuaq basin are excellently exposed but largely inaccessible. however, they can be photographed from boat or helicopter. the key data for the five geological profiles consists of long series of colour stereo photographs taken with ordinary, but calibrated, small-frame cameras. these photographs have been orientated and measured using multi-model photogrammetry as described by dueholm (1992) and pedersen & dueholm (1992). in combination with field logging, sampling, geochemistry and petrography, this has enabled us to establish a detailed stratigraphical framework for the volcanic rocks and hence to trace the evolution and facies changes of the volcanic and sedimentary units with time. the profiles the individual profiles (fig. 1), at scale 1:20 000, are between 80 km and 131 km long and show altitudes up to 2100 m with no vertical exaggeration. three examples illustrating aspects of the contents and details of the profiles are shown in figs 2-4. figure 2 shows progradation of hyaloclastite breccias and associated lava flows into a marine basin; this structure, with giant foresets in the hyaloclastite breccias, has been widely used as a model for the interpretation of seismic sections in offshore areas. an unusual feature of central nuussuaq is the presence of submarine feeder systems and eruption sites for several of the volcanic units, which make the bedding structures in the hyaloclastites very irregular. the concentration of feeder sites in this small area is due to the presence of a deep fault zone, along which magma ascent from the mantle was focused. figure 3 illustrates late tectonic movements and subsequent dyke intrusion. near the west coast of disko the lava pile is faulted and tilted seawards at low angles that can be read very accurately on the profile. the reading of the dips requires that the strikes are perpendicular to the section which is mostly, but not always, the case; the index maps accompanying each profile sheet show strikes and dips of the lavas along the profiles. figure 4 illustrates sediments and lava flows and their interaction on the flood plain in the eastern part of the basin in north-east disko. fluvial sandstones are interbedded with 54 w e 20 km18 km 23 km point 1613 m point 1350 milugissoq fig. 2. part of the geological section through the aaffarsuaq valley (profile 2), central nuussuaq (pedersen et al. 2002a) at 17-24 km on the section base line. no vertical exaggeration. sediments of the cretaceous atane formation (k1) and the paleocene eqalulik formation (t1) are overlain by several volcanic units of the vaigat formation. the atane formation was faulted and tilted prior to the deposition of the volcanic rocks, and a large fault crosses the profile at 18.5 km west of which the atane formation is below exposure level. during the deposition of the volcanic rocks, subaqueous hyaloclastite breccias and subaerial lava flows prograded into the basin from the west to the east. lava flows and hyaloclastite breccias from the same volcanic unit are shown with respectively darker and lighter colour shades; such subaerial–subaqueous pairs are c2 and c1(a+b); d2 and d1(a+b) (nuuk killeq member), and e2 and e1(a+b) (naujánguit member). the hyaloclastite breccias show east-dipping foreset bedding (d1b) or irregular bedding structures (d1a, e1b) near eruption sites (feeder dykes lined with red). the red unit f1, at around 23-24 km, comprises the pyroclastic rocks of the ilugissoq volcano (asuk member) which erupted on the margin of the shallow marine shelf. the younger units are banked up against the volcanic edifice, and the lavas of the unit i2 (ordlingassoq member) overran and covered it. two units of paleocene black shales, both of which were deposited in large and deep lakes. the paleocene sediments are contemporaneous with volcanic rocks in the western part of the basin; the first volcanic unit to reach so far east was the middle part of the rinks dal member of the maligât formation. the boundaries between the major volcanic units form important near-isochronous surfaces that allow the tracing of regional synand post-volcanic tectonic movements in the nuussuaq basin. for example, the boundary between the upper rinks dal member and the nordfjord member of the maligât formation is tilted below sea level in western disko (fig. 3) and rises eastwards towards the disko gneiss ridge to more than 1600 m altitude in central disko. east of the gneiss ridge the surface is gently tilted towards the south-east by slow synand post-volcanic regional movements, so that in eastern disko (fig. 4) it is situated at 940 m altitude. in eastern nuussuaq the same surface is situated at up to 1800 m altitude (not shown here). 55 11 km 13 kminngissuaq point 639 m nw se fig. 3. part of the geological section along the south-west coast of disko (pedersen et al. 2003) at 10.6-14 km on the section base line (profile 3). no vertical exaggeration. the pile of formerly flat-lying subaerial lava flows is tilted and the lavas dip 2°-4° seawards (to the left), preserving some of the stratigraphically highest lava flows of the maligât formation on the mountain tops. the unit m8.8 is the uppermost part of the rinks dal member in which individual lava flows are indicated with separate colours and annotation (e.g. m8); such individual flows extend typically over 5-15 km and then peter out. units no1 to no2 are crustally contaminated lavas of the nordfjord member, and ni1 to ni4 are crustally contaminated lavas of the niaqussat member, of which ni1 carries native iron (fe). a crater area (cr) with thick, irregular, red-oxidised material is seen in the left part of ni4. the lava pile is cut by younger dykes (δm and δe) of which the eocene δe runs very obliquely to the section. inngigissoq 101 km 103 km t5b nw se fig. 4. part of the geological section across disko, near pingu in eastern disko (pedersen et al. 2005) at 100.6-104 km on the section base line (profile 4). no vertical exaggeration. sediments and lava flows are shown as viewed from the south, but are projected into the section both from the south (upper part) and from the north (lower part below the broken line: exposures along the vaigat coast). fluvial sandstones of cretaceous (k1) and paleocene (t2) age are separated by an interval of uncertain age (t0). two lake complexes each show filling with a coarsening-upwards succession of black mudstones and fluvial sandstones: the pingu member (t3a-t3b) and the assoq member (t5a-t5b). the level marked ’u’ is an unconformity within the sandstones. lava flows invaded the unconsolidated sediments to form three sill-like bodies (invasive lavas, in), and finally the sediments were overrun by entirely subaerial lava flows. the entablature zone (e) of the lowest flow indicates a moist environment. the lava flows are from the maligât formation, upper rinks dal member (m8) and nordfjord member (iron-bearing no1, and no2). perspectives the completion of the five profiles, and the complementary geological maps at scale 1:100 000, has provided a significantly improved understanding of the overall three-dimensional structure and stratigraphy of the southern part of the nuussuaq basin, in particular of the well-exposed volcanic formations. in addition to the published profiles, an extensive web of colour stereo panels has been constructed wherever helicopter logistics allowed the photography, so that the total coverage now exceeds 1500 km of stereo panels. the existence of this web, assisted by complete on-line coverage of vertical aerial photograph models, has opened for a range of quantitative geological studies centred on volcanology, sedimentology, palaeomagnetics and tectonics. using palaeomagnetic information, pedersen et al. (2002b) calculated eruption rates and basin subsidence rates for an early part of the vaigat formation, and found eruption rates to be much in excess of those characterising modern iceland. measurements of lake depths and lateral correlation of subaqueous volcanic units over more than 20 km were instrumental for the understanding of the environment that gave rise to unique subaqueous rootless volcanic cones in the assoq lake (larsen et al. in press). the ability of the photogrammetrical method to identify volcanic eruption sites within the web of stereo panels has been demonstrated by the discovery of a large number of crater sites within the vaigat formation on nuussuaq. of particular significance is the observation that economically interesting crustally contaminated volcanic rocks tend to erupt along older fault zones within the basin. such an example is provided by the ilugissoq graphite andesite volcano (pedersen & larsen in press), which is the largest eruption site within the vaigat formation. it is believed that continued 3-d based work within the nuussuaq basin will be of considerable interest for a broad range of geoscientists from both universities and industry. acknowledgements the work was partly supported by the carlsberg foundation, grant no. 0385/30. references chalmers, j.a., pulvertaft, t.c.r., marcussen, c. & pedersen, a.k. 1999: new insight into the structure of the nuussuaq basin, central west greenland. marine and petroleum geology 16, 197–224. dueholm, k.s. 1992: geological photogrammetry using standard smallframe cameras. in: dueholm, k.s. & pedersen, a.k. (eds): geological analysis and mapping using multi-model photogrammetry. rapport grønlands geologiske undersøgelse 156, 7–17. larsen, l.m., pedersen, a.k. & pedersen, g.k. 2006: a subaqueous rootless cone field at niuluut, disko, paleocene of west greenland. lithos 92, 20–32. pedersen, a.k. & dueholm, k.s. 1992: new methods for the geological analysis of tertiary volcanic formations on nuussuaq and disko, central west greenland, using multi-model photogrammetry. in: dueholm, k.s. & pedersen, a.k. (eds): geological analysis and mapping using multi-model photogrammetry. rapport grønlands geologiske undersøgelse 156, 19–34. pedersen, a.k. & larsen, l.m. 2006: the ilugissoq graphite andesite volcano, nuussuaq, central west greenland. lithos 92, 1–19. pedersen, a.k., larsen, l.m. & dueholm, k.s. 1993: geological section along the south coast of nuussuaq, central west greenland. 1:20 000 coloured geological sheet. copenhagen: geological survey of greenland. pedersen, a.k., larsen, l.m. & dueholm, k.s. 2002a: geological section along the north side of the aaffarsuaq valley and central nuussuaq, central west greenland. 1:20 000 coloured geological sheet. copenhagen: geological survey of denmark and greenland. pedersen, a.k., larsen, l.m., riisager, p. & dueholm, k.s. 2002b: rates of volcanic deposition, facies changes and movements in a dynamic basin: the nuussuaq basin, west greenland, around the c27n–c26r transition. in: jolley, d.w & bell, b.r. (eds): the north atlantic igneous province: stratigraphy, tectonics, volcanic and magmatic processes. geological society (london) special publications 197, 157–181. pedersen, a.k., larsen, l.m., pedersen, g.k., heinesen, m.v. & dueholm, k.s. 2003: geological section along the south and south-west coast of disko, central west greenland. 1:20 000 coloured geological sheet. copenhagen: geological survey of denmark and greenland. pedersen, a.k., larsen, l.m., pedersen, g.k. & dueholm, k.s. 2005: geological section across north central disko from nordfjord to pingu, central west greenland. 1:20 000 coloured geological sheet. copenhagen: geological survey of denmark and greenland. pedersen, a.k., larsen, l.m., pedersen, g.k., sønderholm, m., midtgaard, h.h., pulvertaft, t.c.r. & dueholm, k.s. 2006: geological section along the north coast of the nuussuaq peninsula, central west greenland. 1:20 000 coloured geological sheet. copenhagen: geological survey of denmark and greenland. storey, m., duncan, r.a., pedersen, a.k., larsen, l.m. & larsen, h.c. 1998: 40ar/39ar geochronology of the west greenland tertiary volcanic province. earth and planetary science letters 160, 568–586. 56 authors’ addresses a.k.p., geological museum, university of copenhagen, øster voldgade 5–7, dk-1350 copenhagen k, denmark. e-mail: akp@snm.ku.dk l.m.l., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. g.k.p., geological institute, university of copenhagen, øster voldgade 10, dk-1350 copenhagen k, denmark. k.s.d., information and mathematical modelling, richard petersens plads, dtu-bygning 321, dk-2800 lyngby, denmark. geological survey of denmark and greenland bulletin 18, 2009, pp. 126 + map (not online) geological survey of denmark and greenland bulletin 18 � 2009 greenland from archaean to quaternary descriptive text to the 1995 geological map of greenland, 1:2 500 000. 2nd edition niels henriksen, a.k. higgins, feiko kalsbeek and t. christopher r. pulvertaft geological survey of denmark and greenland ministry of climate and energy geological survey of denmark and greenland bulletin 18 keywords archaean, caledonides, cenozoic, economic geology, geological map, greenland, ice sheet, mesozoic, offshore, orogenic belts, palaeozoic, petroleum, phanerozoic, proterozoic, sedimentary basins. cover illustration the cover design depicts mountains of the east greenland caledonian fold belt. the view, west of mestersvig (located on map, page 4), is north over bersærkerbræ and the northern part of the stauning alper to kong oscar fjord with traill ø in the right background. the mountains up to 1800 m high are of the neoproterozoic eleonore bay supergroup. to the right: first author niels henriksen, for many years head of geological mapping at ggu/geus, and participant in field work in greenland for more than 45 years. frontispiece: facing page major caledonian syncline involving reactivated archaean basement gneisses containing amphibolite bands. overlying rusty coloured mesoproterozoic metasediments (krummedal supracrustal sequence) just visible in tight core of the fold. the intensity of deformation in the syncline clearly increases towards the core, where the basement gneisses become more strongly foliated. some of the amphibolite bands were derived from cross-cutting basic intrusions, which are still discernable in the less severely deformed parts of the archaean basement (fig. 17, p. 31). the height of the section is c. 2000 m. south-west of innermost nordvestfjord / kangersik kiatteq (c. 71°30´n), scoresby sund region, central east greenland. chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, geological institute, university of aarhus; minik rosing, geological museum, university of copenhagen; finn surlyk, department of geography and geology, university of copenhagen scientific editor: adam a. garde editorial secretaries: jane holst and esben w. glendal referees: c.r.l. friend (uk), gordon n. oakey (canada) and henrik stendal (greenland) illustrations: eva melskens, helle zetterwall, lis duegaard and jette halskov digital photographic work: benny m. schark graphic production: annabeth andersen printers: rosendahls schultz grafisk, albertslund, denmark manuscript submitted: 29 may, 2009 final version approved: 8 october 2009 printed: 27 november 2009 isbn 987-87-7871-258-5 issn 1604-8156 citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 18, 126 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk or at www.geus.dk/publications/bull © de nationale geologiske undersøgelser for danmark og grønland (geus), 2009 for the full text of the geus copyright clause, please refer to www.geus.dk/publications/bull canada station nord ellesmere island peary land western eastern central daneborg mestersvig pituffik / thule air base melville bugt svartenhuk halvø disko disko bugt upernavik nuuk aasiaat kangerlussuaqsisimiut tasiilaq / ammassalik scoresby sund blossevil le k ys t ivittuut qaqortoq kap farvel nare s s trai t central east southern east north east iceland north west north cent ral west southern west south east south west south arctic ocean north atlantic ocean inland ice 400 km 72° 76° 80° 28°36°44°52° 68° 64° 60° geographical subdivisions of greenland used by the survey map showing the survey’s geographical subdivisions of greenland, both onshore and offshore, used in this bulletin. thus nares strait, the seaway separating greenland from ellesmere island, canada, borders north-west greenland and western north greenland. however, application of these subdivisions has not been rigorously applied in all parts of the text in order to avoid unwieldy phrases. it should also be noted that ‘west greenland’ and ‘east greenland’ are used both as shown on the map (each with two subdivisions, viz. central and southern) and for the entire western and eastern sides of greenland. in this broader sense, west greenland covers the four western subdivisions, viz. north-west, central west, southern west and south-west greenland whereas east greenland covers the four eastern subdivisions, viz. north-east, central east, southern east and south-east greenland. the subdivisions are used throughout the text and also in the legend explanation (pp. 110–112) and index (pp. 117 –126). 4 editorial note this bulletin is a revised 2nd edition of the descriptive text to the geological map of greenland 1:2 500 000. the description was first published in 2000 as geology of greenland survey bulletin 185 by henriksen et al. 2000. the map, compiled by escher & pulvertaft in 1995 and printed the same year, accompanies the present description. since the first edition of this work was published, large amounts of new data have been acquired, especially in the offshore regions, in relation to mineral prospecting and in connection with general geological research mainly in west greenland. the present description aims at providing an up dated overview of the geology of greenland with reference to the enclosed geolo gical map from 1995 that in general terms is still valid. the first edition included an extensive reference list designed as a key to the most relevant sources for the explanation of the geological map of greenland 1:2 500 000. in this second edition the reference list has been expanded with more than 200 new references to cited papers, to give the reader a possibility to follow up on new data and details in agreement with modern interpretations. the 1:2 500 000 map presents a general overview of green land ge ology, but as a basis for this overview there also exists a wealth of more detailed published maps. a set of 14 geological maps at scale 1:500 000 covers the onshore areas of the entire country and a special map at scale 1:1 000 000 covers onshore areas in north-east greenland. in addition to these maps more than 60 geological maps at 1:100 000 have been published, covering mainly areas in central and southern west greenland. a wide range of special geological and geophysical maps has also been published covering both onshore and offshore areas. details of all these publications can be obtained from the survey’s website. for a catalogue of greenland publications and data see: www.geus.dk/publications/publ-uk.htm. this list will be updated at intervals, when relevant new data become available. 5 abstract . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . preface . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . introduction. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . concept of the geological legend . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . topographic base . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . crystalline rocks older than 1600 ma: the greenland precambrian shield . . . . . . . . . . archaean craton . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . eoarchaean supracrustal rocks. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . eoarchaean (‘amîtsoq’) gneisses . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . mesoand neoarchaean supracrustal rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . anorthositic rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . mesoand neoarchaean gneisses . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . intrusive rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . palaeoproterozoic orogenic terrains . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . nagssugtoqidian orogen, west greenland. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . supracrustal rocks. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . felsic and intermediate intrusions. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . nagssugtoqidian orogen in south-east greenland . . . . . . . . . . . . . . . . . . . . . . . . . palaeoproterozoic plutonic rocks. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . rinkian fold belt. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . archaean and palaeoproterozoic supracrustal rocks: disko bugt and nuussuaq . gneisses and intrusions: disko bugt and nuussuaq . . . . . . . . . . . . . . . . . . . . . . palaeoproterozoic supracrustal rocks north of nuussuaq: the karrat group . . . . gneisses and intrusive rocks north of nuussuaq . . . . . . . . . . . . . . . . . . . . . . . . . north-west greenland and the inglefield orogenic belt . . . . . . . . . . . . . . . . . . . . . ketilidian orogen . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . palaeoproterozoic supracrustal rocks in the northern border zone . . . . . . . . . . . . palaeoproterozoic granitoids and basic–intermediate intrusions, the julianehåb batholith . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . metasedimentary rocks in the south-eastern part of the ketilidian orogen . . . . . the ketilidian rapakivi suite. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . archaean–palaeoproterozoic basement in the east greenland caledonian orogen. . archaean–palaeoproterozoic basement beneath the inland ice . . . . . . . . . . . . . . . . proterozoic to phanerozoic geological development after formation of the precambrian shield palaeoto mesoproterozoic unfolded units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . independence fjord group, north greenland . . . . . . . . . . . . . . . . . . . . . . . . . . . . zig-zag dal basalt formation, north greenland . . . . . . . . . . . . . . . . . . . . . . . . . . correlation with similar rocks in the northernmost part of the east greenland caledonides . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . gardar province, south greenland . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . early neoproterozoic orogenic units reworked in the east greenland caledonian fold belt . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . supracrustal rocks. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . migmatites and granites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . mesoproterozoic – early neoproterozoic sedimentary basin in north-west greenland and ellesmere island . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9 11 12 12 13 14 15 17 18 19 19 19 20 21 22 23 24 24 24 25 26 26 26 27 27 28 28 29 30 30 31 32 33 33 33 33 35 35 38 38 39 39 contents 6 thule supergroup . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . neoproterozoic sedimentary basins in north, north-east and east greenland . . . . . hagen fjord group, north greenland. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . eleonore bay supergroup, east and north-east greenland . . . . . . . . . . . . . . . . . tillite group, east greenland . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . sedimentary rocks of unknown age in the east greenland caledonides . . . . . . . . . . carbonatites, kimberlites and associated rocks, west greenland. . . . . . . . . . . . . . . . the palaeozoic franklinian basin of north greenland and ellesmere island . . . . . . uppermost neoproterozoic – silurian in north greenland . . . . . . . . . . . . . . . . . proterozoic–silurian exotic terrane of ellesmere island (pearya) . . . . . . . . . . . . . . ellesmerian orogeny in north greenland and ellesmere island . . . . . . . . . . . . . . . . lower palaeozoic of east greenland . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . cambrian–ordovician sediments in the caledonian fold belt . . . . . . . . . . . . . . . caledonian orogeny in east and north-east greenland . . . . . . . . . . . . . . . . . . . . . . caledonian intrusions and plutonic rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . devonian continental sediments in east greenland . . . . . . . . . . . . . . . . . . . . . . . . . carboniferous–tertiary deposits of the wandel sea basin, central and eastern north greenland . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . late palaeozoic and mesozoic rift basins in east greenland . . . . . . . . . . . . . . . . . . . late permian – early cretaceous deposits of the jameson land basin (70–72°n) late permian – cretaceous deposits in north-east greenland (72–76°n) . . . . . . cretaceous–palaeogene deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . central west greenland . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . southern east greenland . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . tertiary volcanics, intrusions and post-basaltic sedimentary rocks . . . . . . . . . . . . . . palaeogene basalts, central west greenland . . . . . . . . . . . . . . . . . . . . . . . . . . . . . palaeogene basalts, east greenland. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . blosseville kyst region (68–70°n) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . hold with hope to shannon region (73–75°n). . . . . . . . . . . . . . . . . . . . . . . . palaeogene intrusions, east greenland . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . post-basaltic palaeogene sedimentary rocks, east greenland . . . . . . . . . . . . . . . . . pliocene–pleistocene sediments, central north greenland . . . . . . . . . . . . . . . . . . . . quaternary glacial sediments and glaciation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . glaciology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . mass balance . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . past climate and environment. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . offshore geology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . the continental margin off east and north greenland . . . . . . . . . . . . . . . . . . . . . . east greenland south of 77°n . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . north-east greenland (77–83°n) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . the morris jesup rise and the yermak plateau . . . . . . . . . . . . . . . . . . . . . . . . . . the continental margin off west greenland. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . southern west greenland (58–64°n) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . davis stræde (c. 64–69°n). . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . baffin bugt (69–77°n) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . nares stræde (77–82°n) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . the continental margin off north greenland, west of the morris jesup rise . . . . . . offshore sedimentary basins. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . north-east greenland shelf (72–80°n) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . liverpool land basin, central east greenland (69°30´–72°n) . . . . . . . . . . . . . . . 39 40 40 41 42 43 44 45 45 48 49 50 50 50 51 52 54 55 55 57 57 57 60 60 60 61 61 61 62 62 62 63 64 64 65 66 66 66 69 70 70 70 71 72 72 73 73 73 75 blosseville kyst basin, east greenland (67–69°30´n). . . . . . . . . . . . . . . . . . . . . . . southern west greenland (60–68°n). . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . central west greenland (68–73°n) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . north-west greenland (73–77°n) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . offshore north greenland; the lincoln sea basin . . . . . . . . . . . . . . . . . . . . . . . . . mineral deposits. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . mineral occurrences in specific geological settings. . . . . . . . . . . . . . . . . . . . . . . . . . . . archaean–palaeoproterozoic high-grade regions . . . . . . . . . . . . . . . . . . . . . . . . . . . mesoproterozoic intracratonic intrusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . palaeozoic orogenic belts. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . upper palaeozoic – mesozoic basins . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . late phanerozoic intrusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . petroleum potential of greenland . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . onshore basins . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . franklinian basin, north greenland (80–83°n). . . . . . . . . . . . . . . . . . . . . . . . . . . late palaeozoic – mesozoic basins, eastern north greenland (80–83°n) . . . . . . . . late palaeozoic – mesozoic rift basins, north-east greenland (72–76°n) . . . . . . . jameson land basin, central east greenland (70°30´–72°n) . . . . . . . . . . . . . . . . . cretaceous–palaeogene basin, central west greenland (69–72°n) . . . . . . . . . . . . . offshore basins . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . north-east greenland shelf (75–80°n) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . liverpool land basin, central east greenland (69°30´–72°n) . . . . . . . . . . . . . . . . blosseville kyst basin, east greenland (67–69°30´n). . . . . . . . . . . . . . . . . . . . . . . west greenland . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . south and south-west greenland (c. 57–62°n) . . . . . . . . . . . . . . . . . . . . . . . . . . . southern west greenland (c. 62–68°n) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . central west greenland (c. 68–73°n) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . north-west greenland (73–77°n) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . western north greenland (north of 80°n). . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . references . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . legend explanation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . place names register . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . index . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . enclosure: geological map of greenland, 1:2 500 000 . . . . . . . . . . . . . . . . . . . . . 7 75 75 76 76 77 78 80 80 82 82 83 83 84 84 84 84 86 86 86 87 87 87 88 88 88 88 89 90 90 91 92 110 113 117 in pocket 8 9 henriksen, n., higgins, a.k., kalsbeek, f. & pulvertaft, t.c.r. 2009: greenland from archaean to quaternary. descriptive text to the 1995 geological map of greenland, 1:2 500 000. 2nd edition. geological survey of denmark and greenland bulletin 18, 126 pp. + map. the geological development of greenland spans a period of nearly 4 ga, from eoarchaean to the quaternary. greenland is the largest island on earth with a total area of 2 166 000 km2, but only c. 410 000 km2 are exposed bedrock, the remaining part being covered by a major ice sheet (the inland ice) reaching over 3 km in thickness. the adjacent offshore areas underlain by continental crust have an area of c. 825 000 km2. greenland is dominated by crystalline rocks of the precambrian shield, which formed during a succession of archaean and palaeoproterozoic orogenic events and stabilised as a part of the laurentian shield about 1600 ma ago. the shield area can be divided into three distinct types of basement provinces: (1) archaean rocks (3200–2600 ma old, with local older units up to >3800 ma) that were almost unaffected by proterozoic or later orogenic activity; (2) archaean terrains reworked during the palaeoproterozoic around 1900–1750 ma ago; and (3) terrains mainly composed of juvenile palaeoproterozoic rocks (2000–1750 ma in age). subsequent geological developments mainly took place along the margins of the shield. during the proterozoic and throughout the phanerozoic major sedimentary basins formed, notably in north and north-east greenland, in which sedimentary successions locally reaching 18 km in thickness were deposited. palaeozoic orogenic activity affected parts of these successions in the ellesmerian fold belt of north greenland and the east greenland caledonides; the latter also incorporates reworked precambrian crystalline basement complexes. late palaeozoic and mesozoic sedimentary basins developed along the continent–ocean margins in north, east and west greenland and are now preserved both onshore and offshore. their development was closely related to continental break-up with formation of rift basins. initial rifting in east greenland in latest devonian to earliest carboniferous time and succeeding phases culminated with the opening of the north atlantic ocean in the late paleocene. sea-floor spreading was accompanied by extrusion of palaeogene (early tertiary) plateau basalts in both central west and central–southern east greenland. during the quaternary greenland was almost completely covered by ice, and the present day inland ice is a relic from the pleistocene ice ages. vast amounts of glacially eroded detritus were deposited on the continental shelves around greenland. mineral exploitation in greenland has so far encompassed cryolite, lead-zinc, gold, olivine and coal. current prospecting activities in greenland are concentrated on gold, base metals, platinumgroup elements, molybdenum, iron ore, diamonds and lead-zinc. hydrocarbon potential is confined to the major phanerozoic sedimentary basins, notably the large basins offshore north-east and west greenland. while reserves of oil or gas have yet to be found, geophysical data com bined with discoveries of oil seeps onshore have revealed a considerable potential for offshore oil and gas. abstract authors’ address geological survey of denmark and greenland, øster voldgade 10, dk 1350 copenhagen k, denmark. e-mail: nielshen@mail.dk 10 6 9 3 85 211 6 9 3 10 47 47 10 1313 1212 85 211 fig. 1. index map of the geological map of greenland, 1:2 500 000, showing segments 2–13, so numbered in the atlas version of the map (for atlas format, see p. 11). segment 1 – the title page of the atlas – is not shown. 11 greenland is the largest island on earth with a surface area of more than two million square kilometres. it is up to 1250 km from east to west and 2675 km from north to south, extending over almost 24 degrees of latitude; the northern extremity is the northernmost land area in the world. the inland ice, the large central ice sheet which covers about 1 756 000 km2 (c. 81%) of greenland, has a maximum thickness of c. 3.4 km. the ice-free strip of land surrounding the inland ice, in places up to 300 km wide, has an area of c. 410 000 km2; this is approximately 30% more than that of the british isles. this ice-free zone is generally very well exposed and yields a wealth of geological information, notably in fjord walls and in mountainous areas; even lowland areas only have a limited vegetation cover due to the arctic setting. the area of greenland’s continental shelf that is underlain by continental crust is estimated to be approximately 825 000 km2. geological observations in greenland began with the first scientific expeditions; these reached west and east green land in the early 1800s and north greenland in the late 1800s and early 1900s. systematic geological mapping commenced in east greenland with lauge koch’s ‘danish expeditions to east greenland’, which lasted from 1926 until 1958 and were mainly concentrated in the region 72–76°n. in west greenland systematic geological investigations began in 1946 with the foundation of grønlands geologiske undersøgelse (ggu – the geological survey of greenland); work was initially concentrated in west greenland but was subsequently extended to all parts of greenland. comprehensive investigations by ggu expanded to include not only geological mapping, but a wide range of geochemical, geophysical and glaciological studies both onshore and offshore. in 1995 ggu was merged with danmarks geologiske undersøgelse (dgu – the geological survey of denmark) to form the present survey, de nationale geologiske undersøgelser for danmark og grønland (geus – the geological survey of denmark and greenland). the broad range of geological activities in greenland previously undertaken by ggu continues to be carried out by geus. when ggu published the first general geological map of all of greenland at scale 1:2 500 000 in 1970, representation of the geology was restricted to onshore areas; relatively little was then known of the offshore geology. during the past almost 40 years the offshore areas surrounding greenland have been investigated by airborne and ship borne geophysical surveys operated by the survey, by other scientific institutions and by commercial companies. when the new edition of the geological map was printed in 1995 enough was known to enable an interpretation of the offshore geology to be presented on the map, although it was emphasised that for some of the remote areas offshore north greenland knowledge was very limited. since 1995 a considerable amount of new information has been collected from the offshore areas. the new knowledge is reported in the chapter on offshore geology in the present description. there is a high petroleum exploration interest in many of the offshore sedimentary basins, and geological knowledge of the offshore areas has been consid er ably augmented as a result of commercial exploration. the geological map of greenland, 1:2 500 000, printed in 1995, is available in three formats: 1. a wall map (sheet size 96 × 120 cm). 2. a folded map sheet (24 × 20 cm, as in the pocket of this bulletin). 3. an atlas of numbered segments (24 × 20 cm when closed. an index to the 12 numbered segments of the map is shown as figure 1. the description of the map has been prepared with the needs of the professional geologist in mind; it requires knowledge of geological principles but not previous knowledge of greenland geology. throughout the text, reference is made to the key numbers in the map legend representing geological units and indicated in square brackets [ ] (see legend explanation, p. 110), while a place names register (p. 113) and an index (p. 117) include place names, geological topics, stratigraphic terms and units found in the legend. the extensive reference list is intended as a key to the most relevant information sources. the text has been compiled by n. henriksen. principal contributors include: n. henriksen (several sections and illustrations); a.k. higgins (neoproterozoic – lower palaeozoic in north greenland and palaeozoic fold belts in north and north-east green land); f. kalsbeek (precambrian shield and palaeo– mesoproterozoic deposits in north greenland); t.c.r. pulvertaft (offshore geology). chapters which have been much revised have been externally reviewed (precambrian shield by c.r.l. friend; offshore geology by g.n. oakey and mineral deposits and petroleum potential by h. stendal). preface 12 a general overview of the geology of the whole of greenland in the form of a coloured map sheet, the tectonic/geological map of greenland, was published by the geological survey of greenland (ggu) in 1970 at a scale of 1:2 500 000. subsequently, much new information became available as a result of new systematic geol ogical mapping in the ice-free land areas, notably in north, north-east and south-east greenland, while offshore areas were investigated by a series of seismic, gravimetric and aeromagnetic surveys. the inland ice was also further investigated in the period 1970–1995 by regional satellite and airborne radar surveys as well as by ground studies and deep drilling through the more than 3000 m thick central part of the ice sheet. at the time of compilation in 1995 the map therefore included much new information from the ice-free land areas, and the offshore regions were represented for the first time on a map at this scale. for the inland ice new representations of the upper and lower surface of the ice sheet were shown by contours, together with its calculated thickness. however, between 1995 and 2009 much additional information has been collected especially from the offshore regions, and as a consequence the enclosed geological map of greenland, 1:2 500 000 does not represent a fully up-to-date expression of the present knowledge on the geology. the 1995-depiction of the ice-free land areas is in general terms still satisfactory, as well as the representation of the inland ice. it is mainly in the offshore regions that new information has been added since 1995, and these developments are described in the chapters on offshore geology and petroleum potential of greenland. another field with new developments is that of mineral occurrences, which are addressed in the chapter on mineral deposits. in order to relate the geology of greenland to neighbouring countries within the borders of the map sheet, the geology of the adjacent areas of canada and iceland has been included, based on published maps (see map legend). the geology of the ice-free land areas on the 1:2 500 000 scale map has been compiled as a conventional bedrock geological map, together with representations of the major tectonic features in the orogenic belts. the presentation of the geology of offshore areas follows a different concept, as interpretations are based on geophysical information. onshore superficial deposits of quaternary age have been shown only where extensive areas of bedrock are covered. in many regions dykes are a prominent element of the geology, but as they only form a minor proportion of the exposures they cannot generally be represented at the scale of the map. a compilation of dykes of different ages is shown in this volume as fig. 20 (p. 37). the inland ice and the many local ice caps and glaciers are shown as one unit. the sea ice, which covers substantial parts of the oceans bordering north and east greenland for much of the year, is not depicted on the map. the term ‘tertiary’ and division of the proterozoic into early, middle and late, was used in the legend on the printed map in accordance with the standards of the early 1990s. in the present revised edition of the map description the current conventions of palaeogene/neogene and palaeo-, mesoand neoproterozoic are introduced. in the precambrian descriptions the prefixes ‘early’, ‘middle’ and ‘late’ have been modernised to the subdivisions eo, palaeo, meso and neo for both archaean and proterozoic time and rock units. concept of the geological legend two different legend concepts have been used – one for the onshore ice-free areas and one for the offshore regions. in the legend for the ice-free land areas a distinction has been made between rocks older and younger than 1600 ma. in the older group, which mainly comprises crystalline rocks of the stable precambrian greenland shield, the rock units are distinguished according to their lithology and age; the extent of regional tectono-metamorphic provinces is also depicted. rocks younger than 1600 ma are shown in relation to the formation of sedimentary basins and orogenic belts along the margins of the stable shield. the principal subdivisions depicted on the map illustrate the general depositional environment, age and extent of the main sedimentary and volcanic basins and, in the franklinian basin in north greenland, the overall depositional setting. younger crystalline gneis ses and plutonic rocks are distinguished by lithology and age of orogenic formation and emplacement. a schematic chronological representation of the geological units shown on the map forms the basic division of the map legend. introduction the structures and the ages of deformation in the various orogenic belts are shown by structural trend lines and major tectonic features by appropriate symbols. most orogenic belts are of composite origin and may incorporate older crystalline rocks and structures. it is often difficult or impossible to distinguish between the older and younger structural elements, and therefore only the signature for the youngest orogenic event has been used within a specific fold belt. post-orogenic undeformed rocks can be recognised by the absence of overprints of structural symbols. a cartoon of the crustal evolution of greenland is shown above the legend. six stages of evolution are shown from the eoarchaean to the cenozoic. these show the distribution in time and space of the orogenic belts and the stepwise growth of the stable crust. the post-orogenic development of sedimentary basins and volcanic provinces is also shown, together with the approximate extent of continental crust around greenland. the legend concept for the offshore areas was based on geological interpretation of the available geophysical data. distinction is made between areas underlain by continental and oceanic crust, respectively; a transition zone is also recognised. areas with oceanic crust are further subdivided into time slices of 15 ma based on magnetic anomaly patterns. magnetic anomaly lines with chron numbers are shown, together with spreading axes and transform faults. major sedimentary basins are indicated by isopachs showing the sediment thickness superimposed on a representation of crustal type. volcanic rocks exposed on the seabed (mostly palaeogene in age) are also shown. an updated overview map of the offshore regions is shown on pp. 68 and 85. topographic base the topographic base for the 1:2 500 000 geological map has been drawn on the basis of fixed points established throughout greenland by kort & matrikelstyrelsen, denmark (kms – the national survey and cadastre, which incorporates the former geodetic institute). the map is constructed as a utm projection in zone 24 with wgs 84 datum; the central meridian is 39°w. photo grammetric constructions by kms and ggu have been combined and co-ordinated to produce the first geometrically correct topographic representation of all of greenland. all previous maps have suffered to varying degrees from insufficient ground control, especially in north greenland where errors in the location of topographic features of up to 25 km occur on older maps. height contours have been omitted on the ice-free land areas to avoid obscuring the geological detail, but they are shown on the inland ice. place names are indicated in both their greenlandic and danish forms, the greenlandic names with the new orthography as approved by the greenland place names authority. a register of place names used on the map is given on p. 113. the bathymetry of the offshore areas has been compiled from various sources. the available material is very heterogeneous, ranging from very detailed navigation maps by the royal danish hydrographic office (now part of kms) to generalised small-scale international oceanographic maps. information from the ice-covered regions off north and east greenland is limited; hydrographic representations from these areas should therefore be viewed with reservation. a topographic map of greenland at a scale of 1:2 500 000 was published by kms in 1994 (kms 1994). the enclosed geological map of greenland at the same scale uses an identical topographic base map with the same projection; the only significant topographical difference is the omission of contour lines on the land areas. based on the digital data for the topographic map, the size of greenland and its ice cover has been computed by weng (1995). the area figures are: ice-free land area 410 449 km2 ice-covered area 1 755 637 km2 total area 2 166 086 km2 13 14 about half of the ice-free area of greenland consists of archaean and palaeoproterozoic crystalline basement rocks, mainly orthogneisses with enclaves of supracrustal rocks. they belong to three distinct kinds of basement provinces (fig. 2): (1) archaean rocks (3200–2600 ma with local older units, up to >3800 ma in the godt håbsfjord region), strongly deformed during the archaean but almost unaffected by proterozoic or later orogenic activity; (2) archaean terrains reworked during the palaeo proterozoic around 1900–1800 ma ago; (3) terrains mainly composed of juvenile palaeoproterozoic rocks (2000–1750 ma). terrains of categories (2) and (3) often contain high-grade palaeoproterozoic metasedimentary successions. nearly all unreworked archaean gneisses occur within the archaean craton of southern greenland (fig. 2). they are cut by swarms of mafic dykes (see fig. 20), most of which were emplaced between 2200 and 2000 ma ago; these dykes are generally undeformed and unmetamorphosed, demonstrating that the surrounding gneisses cannot have been significantly affected by palaeoproterozoic orogenic activity 1900–1800 ma ago. reworked archaean orthogneisses are prominent in the nagssugtoqidian orogen and the rinkian fold belt north of the archaean craton in west greenland, and in the ammassalik region in south-east greenland (fig. 2). reworked archaean gneisses are also exposed in a small area at victoria fjord in northernmost greenland (c. 3400 ma, nutman et al. 2008a) and similar rocks have been found at a locality beneath the inland ice by drilling (weis et al. 1997). juvenile palaeoproterozoic gneisses and granitoid rocks (2000–1750 ma) make up most of the ketilidian orogen of south greenland and parts of the inglefield orogenic belt in north-west greenland. they also form a large proportion of the crystalline basement within the caledonian orogen of north-east greenland. before the opening of the labrador sea and baffin bugt the precambrian basement of greenland formed an integral part of the laurentian shield. a recent interpretation of the relationships between geological provinces in eastern canada and greenland (st-onge et al. 2009) is shown in fig. 3. crystalline rocks older than 1600 ma: the greenland precambrian shield upernavik inglefield orogenic belt victoria fjord uummannaq disko bugt nuuk / godthåb isukasia ketilidian orogen archaean craton rinkian fold belt nagssugtoqidian orogen c aledonian fold belt inland ice tasiilaq / ammassalik gisp 2 400 km preserved archaean rocks reworked archaean rocks juvenile palaeoproterozoic rocks ++ fig. 2. simplified map showing the distribution of archaean and palaeoproterozoic basement provinces in greenland. large areas within the rinkian fold belt are dominated by metasedimentary rocks ( : karrat group) and granites (+: prøven igneous complex). black dots and open circles indicate localities where the presence of, respectively, archaean and palaeoproterozoic rocks have been documented in poorly known areas, as well as in cases where these ages are in contrast to the age of the surrounding rocks. slightly modified from kalsbeek (1994). 15 archaean craton together with smaller areas along the coast of labrador and in north-western scotland, the archaean rocks of southern greenland (figs 2, 3) form the north atlantic craton. the greenland archaean is largely made up of tonalitic to granodioritic orthogneisses [72, 73], amphibolites [68] and anorthositic rocks [85]. most of these rocks are of mesoto neoarchaean age, 3200–2600 ma, but eoarchaean ortho gneisses [76] and supracrusal rocks [69] (3850–3600 ma) are widely exposed in the godt håbsfjord region. before geochronological data became more widely available, the whole of the greenland archaean craton was envisaged to represent a more or less homogeneous geological entity. detailed field investigations combined with u-pb zircon age determinations, however, have ? franklinian basin karrat group aasiaataasiaat domaindomain raerae cratoncraton bas 1.88–1.865 ga srs c. 1.845 ga bes 1.82–1.795 ga tgs 1.87–1.85 ga nis 1.86–1.84 ga dbs c. 1.88 ga superior craton narsajuaq arc makkovikketilidian orogen c. 1.96–1.91 ga c. 1.89–1.80 ga north atlantic craton aasiaat domain franklinian basin rae craton karrat group e ast g reenland c aledonides gl._fig. nyt_fig.03 bulletin 185_2 26-08-09/2 meta incognitameta incognita microcontinentmicrocontinent meta incognita microcontinent 250 km archaean palaeoproterozoic julianehåb batholith ketilidian–makkovik orogen prøven igneous complex, cumberland batholith, lac lomier complex de pas batholith and narsajuaq arc burwell arc piling, hoare bay and karrat groups lake harbour group ramah and mugford groups cape smith – new quebec orogen inglefield orogenic belt and ellesmeredevon terrane proterozoic basement in east greenland mary river group archaean microcontinental fragments rae craton north atlantic craton superior craton palaeogene magmatic province phanerozoic orogens and basins neoproterozoic–palaeozoic basins grenville orogen gardar igneous province and seal lake group mesoproterozoic plutonic suites palaeoto neoproterozoic basins mesoproterozoic–phanerozoic ice fig. 3. map showing the presently preferred correlation of principal geological units of eastern canada and greenland, shown with greenland in its pre-drift (pre-late cretaceous) position relative to eastern canada, simplified after st-onge et al. (2009). bas: baffin suture; bes: bergeron suture; dbs: disko bugt suture; nis: nordre isortoq steep belt; srs: soper river suture; tgs: tasiuyak gneiss suture. the approximate ages of the different sutures illustrate the progressive accretion of crustal blocks from north to south in greenland during the palaeoproterozoic. for details see st-onge et al. (2009) and papers referred to therein. 16 block boundary transition between lower and upper crustal zones mainly upper crustal zone lower crustal zone retrograde amphibolite facies granulite facies ketilidian orogen proterozoic archaean late-tectonic ttg plutons and granitic rocks orthogneiss metavolcanic belt eoarchaean gneiss post-kinematic norites anorthosite complex prograde amphibolite facies 64° 62° 52° q q ta nagssugtoqidian front ta ta t ketilidian front ? ? ? nuuk kangerlussuaq 64° 51° 50 km julianehåb batholith supracrustal rocks reworked archaean rocks granitic rocks orthogneiss granulite facies anorthosite complex supracrustal rocks eoarchaean gneisses palaeoproterozoic mesoproterozoic gardar igneous province archaean q q d a v i s s t r a i t aasivik isukasia akia færingehavn tasiusarssuaq tre brødre sioraq paamiut neria sermiligaarsuk kapisilik qarliit taserssuat assemblage q ta tuno i n l a n d i c e a b m a n iit s o q f is k e f jo r d g o d t h å b s f jo r d a m e r a lik b e lt s e r m ilik b jø r n e s u n d k v a n e f jo r d iv it t u u t i a b i north atlantic craton fig. 4. maps of the archaean craton, southern west greenland. a: the subdivision in terranes/blocks with distinct geological histories (see text) after friend & nutman (2001). t, târtoq group; i, ilivertalik augen granite; q, qôrqut granite complex; ta, taserssuaq tonalite. b: the subdivision in tilted blocks, each of which has granulite facies rocks in its northern parts and rocks at amphibolite facies in the south after windley & garde (2009). shown that areas with contrasting tectonometamorphic histories (‘terranes’) occur side by side, separated by folded mylonite zones (friend et al. 1987, 1988). in the godt håbsfjord region friend et al. (1987) recognised three such terranes, the færingehavn, tre brødre, and tasiu sarssuaq terranes, each with its characteristic rock association and metamorphic history. the subdivision of the godthåbsfjord region into terranes has since been repeatedly revised and refined as more geochronologic information became available (e.g. friend & nutman 2005; nutman & friend 2007). the validity of the terrane model has locally been verified (crowley 2002) but elsewhere questioned (hanmer et al. 2002). after its introduction in the godthåbsfjord region the terrane model has also been applied to other parts of the archaean craton (friend & nutman 2001), see fig. 4a. because of the large areal extent of the archaean rocks many details of the subdivision of the craton into terranes are still to be clarified. the different terranes are envisaged once to have formed independent crustal blocks (perhaps fragments of an earlier archaean continent) that were amalgamated during the neoarchaean. recently, windley & garde (2009) have subdivided the archaean craton of western greenland into six slightly tilted blocks, separated by shear zones, each of which expose granulite facies rocks in the north and (prograde) amphibolite facies rocks in the south (fig. 4b). rock units interpreted as remnants of island arcs are exposed in supracrustal belts in the low-grade parts of these blocks. rocks retrograded from granulite facies to amphibolite facies commonly occur in the northern and central parts of these blocks. the paper of windley & garde (2009) contains an extensive overview of research carried out in the region. eoarchaean supracrustal rocks the isua supracrustal sequence [69] (3700–3800 ma, moorbath et al. 1973) in the isukasia area (fig. 2) at the head of nuup kangerlua/godthåbsfjord is the most extensive occurrence of eoarchaean supracrustal rocks known on earth. it forms a zone up to 4 km wide and up to c. 35 km long and has been investigated in considerable detail. a recent review of earlier studies together with new data and geological maps at 1:20 000 is presented by nutman & friend (2009). these authors subdivide the isukasia area into two tectonic units (terranes), with rocks up to 3700 ma in the north and >3800 ma in the south. the supracrustal rocks comprise: (1) layered and massive amphibolites, within which pillow structures are locally preserved; (2) metacherts and a major body of banded iron formation (fig. 5); (3) biotitemuscovite schists, some of which preserve graded bedding; (4) units of talc schist, up to 100 m wide, with relics of dunite; (5) layered carbonate and calc-silicate rocks, strongly affected by metasomatic activity and (6) bodies of pale chloritic amphibolite (‘garbenschiefer’) up to 1 km wide, which form c. 25% of the supracrustal belt, and probably represent metasomatically altered metavolcanic rocks. all these rocks have been strongly deformed and metamorphosed at amphibolite facies conditions. geochemical studies of the least altered amphibolites have shown that they have tholeiitic and boninitic compositions, similar to modern basaltic rocks formed in oceanic island arcs (polat et al. 2002; polat & hoffman 2003). outside the isukasia area enclaves of supracrustal rocks, mainly amphibolites of tholeiitic or komatiitic composition, occur as thin units within eoarchaean gneisses. these supracrustal rocks have been collectively 17 fig. 5. eoarchaean banded iron formation consisting of interlayered magnetite (dark layers) and chert (light layers). isua supracrustal sequence, isukasia, inner godthåbsfjord, southern west greenland. pen for scale. photo: a.a. garde. termed the akilia association [69], and are thought to represent remnants of a disrupted greenstone belt (mcgregor & mason 1977). studies of graphite particles in samples of isua metasedimentary rocks have yielded evidence of very early life on earth (rosing 1999). eoarchaean (‘amîtsoq’) gneisses eoarchaean orthogneisses (fig. 6), previously known as amîtsoq gneisses and shown under that name on the geological map [76], occur in an area stretching north-east from nuuk/godthåb to isukasia. they are characterised by the presence of abundant remnants of metamorphosed basic dykes (ameralik dykes, fig. 20; mcgregor 1973). the precursors of the gneisses were formed during a number of distinct intrusive events between c. 3800 and 3600 ma (moorbath et al. 1972; nutman et al. 2004). because of the diversity in age and origin of the eoarchaean rocks nutman et al. (1996) introduced the term itsaq gneiss complex to include all eoarchaean rocks in the godthåbsfjord region. the term amîtsoq gneisses is rarely used in newer publications. two main types of eoarchaean orthogneisses (not differentiated on the geological map) can be recognised: (1) grey, banded to homogeneous tonalitic to granodioritic orthogneisses of calc-alkaline affinity (commonly with secondary pegmatite banding) which form at least 80% of the outcrop. the oldest of these have been dated at c. 3850 ma (for overview see nutman et al. 2004), although these very old dates have been questioned (e.g., whitehouse et al. 1999); (2) microcline augen gneisses with associated subordinate ferrodiorites (c. 3600 ma), which have been referred to as the amîtsoq iron-rich suite (nutman et al. 1984). the latter resemble proterozoic rapakivi granites and were intruded after strong deformation of the surrounding grey banded gneisses. most eoarchaean gneisses are in amphibolite facies, but locally the rocks have been affected by c. 3600 ma granulite facies metamorphism possibly related to emplacement of the amîtsoq iron-rich suite. after compilation of the geological map eoarchaean orthogneisses have also been found north of the godt 18 fig. 6. heterogeneous, polyphase eoarchaean (amîtsoq) gneiss in the central part of northern godthåbsfjord, southern west greenland. fragments of dark homogeneous amphibolite are interpreted as remnants of disrupted mafic (ameralik) dykes. the hammer is c. 45 cm long. photo: a.a. garde. håbsfjord region, at qarliit tasersuat (65°49´n, 50°44´w) and in a larger area east of sukkertoppen iskappe, the aasivik terrane (rosing et al. 2001; fig. 4a). mesoand neoarchaean supracrustal rocks ten to twenty per cent of the archaean craton is made up of a variety of supracrustal rocks [68], mainly amphibolites with subordinate paragneisses (often garnetiferous ± cordierite ± sillimanite) and ultramafic layers and pods. amphibolites represent the oldest rocks recognised within each terrane; primary cover–basement relationships with underlying rocks have not been observed. few reliable age determinations for these rocks are available, but it is evident that they belong to several different age groups. amphibolites locally show well-preserved pillow structures indicating a submarine volcanic origin. intense deformation, however, has generally obliterated all primary structures and produced finely layered amphibolites. more massive amphibolites may represent original basic sills within the volcanic pile. the amphibolites range from andesitic to komatiitic in composition; the majority are chemically similar to low-k tholeiitic basalts. two typical examples of mesoarchaean supracrustal units are: (1) andesitic metavolcanic rocks in the southeastern akia terrane (fig. 4a) that have been dated at 3070 ma. on the basis of field observations and chemical data they have been interpreted as parts of a mesoarchaean island arc (garde 2007a). (2) a several kilometres thick, relatively well preserved succession of supracrustal rocks in the southernmost part of the archaean craton, the târtoq group (t, fig. 4a; higgins 1968; berthelsen & henriksen 1975), which consists of metavolcanic greenschists and metasedimentary rocks. while the youngest detrital zircons in a metasedimentary sample are c. 2840 ma (nutman et al. 2004), greenschists at a different locality are cut by sheets of 2940 ma tonalite (nutman & kalsbeek 1994). apparently the târtoq group consists of supracrustal packages of varying ages, illustrating once more the geological complexity of the archaean craton. metasedimentary rocks occur only locally. geochron ological data on detrital zircons suggest a variety of ages (nutman et al. 2004). while the youngest zircons in most of the investigated samples are c. 2800 ma (in samples from the godthåbs fjord region c. 3000 ma), a sample from hamborgerland, north of maniitsoq/suk kertoppen, has zircons as young as 2720 ma. eoarchaean zircons are rare or absent in all of the investigated samples, supporting the view that the eoarchaean terranes were separated from the other terranes until the archaean craton was united by terrane amalgamation during the neoarchaean. most supracrustal units are complexly folded and, since they form good marker horizons, they have been used to reveal the intricate structure of the enveloping gneiss complexes. anorthositic rocks metamorphosed calcic anorthosites and associated leucogabbroic, gabbroic and ultramafic rocks [85] form one of the most distinctive rock associations in the archaean craton. such rocks are present in all the terranes, and detailed investigation has revealed subtle geochemical variations between anorthosites from different terranes (dymek & owens 2001). anorthositic rocks occur as concordant layers and trains of inclusions within gneisses, and provide some of the best marker horizons for mapping structures on a regional scale. they are generally bord ered by amphibolites into which they are believed to have been intruded. anorthosites and associated rocks are most spectacularly developed in the fiskenæsset area of southern west greenland where they form c. 5% of the total outcrop. here they appear to belong to a single stratiform intrusion, the fiskenæsset complex (myers 1985), which has been dated at c. 2850 ma (ashwal et al. 1989). the main rock types are metamorphosed anorthosites, leucogabbros and gabbros (<10%, 10–35% and 35–65% mafic minerals, respectively), together with minor proportions of ultramafic rocks and chromitite (ghisler 1976). although the rocks are commonly strongly deformed, magmatic structures are preserved in low-strain areas: cumulus textures with plagioclase up to 10 cm in size are common and igneous layering can be observed at many localities. the fiskenæsset complex has undergone complex folding. the earliest major folds were recumbent isoclines; these were refolded by two later fold phases producing structures with steeply inclined axial surfaces (myers 1985). mesoand neoarchaean gneisses most of the archaean craton is composed of mesoar chaean grey orthogneisses [72, 73]. in accordance with the notion that the craton comprises a number of individual terranes, variable ages have been reported (for a detailed overview see windley & garde 2009). in the akia terrane (fig. 4a) of west greenland ages up to 19 20 3220 ma occur (garde 1997), whereas farther south isotopic ages are generally less than 3000 ma. the 2825 ma ikkattoq gneisses in the tre brødre terrane, godthåbsfjord region (friend et al. 2009), are an example of these younger gneisses. the ikkatoq gneisses are mainly of granodioritic composition, with subordinate quartz diorite. sm-nd isotope data indicate that eoarchaean sources played a significant role in their petrogenesis. the igneous precursors of the archaean gneisses were intruded as sub-concordant sheets and larger complexes that penetrated and disrupted (‘exploded’) pre-existing basic metavolcanic units and anorthositic rocks (fig. 7); the gneisses commonly occupy much larger volumes than the older rocks into which they were intruded. individual gneiss sheets range from a few metres to several kilometres in width. it has been suggested that intrusion of granitoid magmas was associated with periods of thrusting (bridgwater et al. 1974). most of the gneisses are tonalitic to granodioritic in composition and form typical ttg (tonalite, trond hjemite, granodiorite) suites. a statistical study in the fiskenæsset area (c. 4000 km2) has shown that such gneisses make up c. 85% of the outcrop area. tonalitic gneisses (k-feldspar <10%) form c. 57%, granodioritic gneisses (k-feldspar 10–20%) c. 9%, and granitic gneisses (k-feldspar >20%) c. 16% of the terrain (kalsbeek 1976). large parts of the craton are occupied by granulite fa cies gneisses [73]. granulite facies metamorphism, however, was not synchronous throughout the area: north of nuuk/godthåb it is 3000–3100 ma (garde 1990; friend & nutman 1994), whereas in the fiskenæsset area it is c. 2800 ma (pidgeon & kalsbeek 1978) and north of maniitsoq/sukkertoppen c. 2750 ma (friend & nutman 1994). commonly the age of granulite facies metamorphism is similar to that of the igneous precursors of the gneisses. in granulite facies terrains hypersthene is most common in amphibolites, whereas in orthogneisses its presence depends on chemical composition. regional surveys of stream sediment geochemistry have shown that the distribution of several lithophile elements is strongly correlated with metamorphic facies variations (fig. 8; steenfelt 1994). two kinds of amphibolite facies gneisses [72] can be distinguished: those formed by retrogression of granulite facies rocks, and those that were formed by prograde metamorphism and never experienced granulite facies conditions. these two kinds have not been differentiated on the geological map because criteria to recognize retrograded granulite facies rocks (mcgregor & friend 1997) were not available during the early mapping. an overview of the distribution of prograde and retrograde amphibolite facies gneisses in southern west greenland is presented by windley & garde (2009) and is here shown in fig. 4b. commonly the gneisses show complex fold interference structures (e.g. berthelsen 1960; fig. 9). formation of gneisses by deformation and migmatisation of their igneous precursors has been described in detail by myers (1978), and a detailed description of the complex evolution of the fiskefjord area, north of godthåbsfjord, has been presented by garde (1997). intrusive rocks within the archaean craton a variety of homogeneous granitic to tonalitic rock units have been differentiated on the map as felsic intrusions [80]. these rocks were em placed at various times during the tectonic evolution of the areas in which they occur. some, e.g. the c. 2980 fig. 7. amphibolite agmatite with numerous sheets of tonalite, granodiorite, granite and pegmatite, dated at 3.0–2.97 ga. south-facing, c. 40 m high cliff in central godthåbsfjord, southern west greenland. person in red anorak for scale. photo: a.a. garde. ma taserssuaq tonalite north of inner godthåbsfjord (ta, fig. 4a; garde 1997), represent late phases of the igneous precursors of the gneisses in areas where deformation was less intense than elsewhere. others, e.g. the 2835 ma ilivertalik augen granite in the tasiusarsuaq terrane (i, fig. 4a; pidgeon et al. 1976) are younger than the surrounding gneisses, but have been strongly overprinted by later deformation and metamorphism. one rock unit, the 2550 ma qôrqut granite complex [79] east of nuuk/godthåb (q, fig. 4a; friend et al. 1985), was formed by late crustal melting and is clearly posttectonic. a distinct 2700 ma suite of very well-preserved posttectonic intermediate and mafic intrusions, including gabbros and diorites [82] as well as syenites and granites [80], occurs within archaean gneisses in the skjoldungen district of south-east greenland (nielsen & rosing 1990; blichert-toft et al. 1995). it is associated with older syenitic gneisses [80] and with a late nephelinite body, the 2670 ma singertât complex [83]. small norite bodies [82] occur within an arcuate belt east of maniitsoq/sukkertoppen (secher 1983), and a small 3007 ma carbonatite sheet [84] (the oldest carbonatite known on earth) has been found at tupertalik, 65°30´n in west greenland (larsen & pedersen 1982; larsen & rex 1992; bizzarro et al. 2002). palaeoproterozoic orogenic terrains about forty per cent of the ice-free area of greenland is underlain by palaeoproterozoic orogenic terrains (fig. 2). north of the archaean craton lies the nagssugtoqidian orogen, which continues beneath the inland ice to south-east greenland. still farther north are the rinkian fold belt and the inglefield orogenic belt of west and north-west greenland, and south of the archaean craton lies the ketilidian orogen (fig. 2). the nagssugtoqidian orogen and rinkian fold belt largely consist of reworked archaean rocks that underwent strong deformation and metamorphism during the palaeoproterozoic 1900–1850 ma ago, while the inglefield and ketilidian orogens contain large proportions of juvenile palaeoproterozoic crust. recent investigations have suggested that the greenland shield was formed by progressive accretion of crustal blocks, from northernmost greenland to the south (nutman et al. 2008b; st-onge et al. 2009). archaean rocks of the rinkian fold belt were united with an archaean block in northernmost greenland along the inglefield orogenic belt around 1920 ma (nutman et al. 2008a). archaean gneisses of the nagssugtoqidian orogen were then accreted 21 45°w45°w 61.761.7 40.140.1 29.529.5 19.019.0 13.613.6 10.410.4 8.28.2 6.46.4 5.25.2 4.04.0 3-13-1 2.12.1 1.71.7 1.11.1 0.80.8 0.10.1 uu ppmppm granulite facies gneiss 50 km 62°n 64°n 66°n 51°w 51°w 45°w 54°w 61.7 40.1 29.5 19.0 13.6 10.4 8.2 6.4 5.2 4.0 3-1 2.1 1.7 1.1 0.8 0.1 u ppm fig. 8. map showing the variation in uranium concentrations in the <0.1 mm fraction of stream sediments from the archaean craton of southern west greenland (steenfelt 2001). there is a close correlation between u concentrations and the metamorphic grade of the underlying rocks (fig. 4 b), stream sediments from granulite facies terrains having low u compared to those from amphibolite grade areas. 22 to the rinkian/inglefield/north greenland block at c. 1870 ma along a suture within disko bugt (connelly & thrane 2005; thrane et al. 2005; connelly et al. 2006). collision within the nagssugtoqidian orogen followed around 1850 ma (connelly et al. 2000), and batholitic rocks of the ketilidian orogen were accreted to the archaean craton 1850–1800 ma ago (garde et al. 2002). many details of this process are still uncertain. the largest area of juvenile palaeoproterozoic rocks in greenland (600 km along strike and up to 300 km in width) occurs in the caledonian thrust sheets of northeast greenland. its relationships with the other palaeo proterozoic terrains in greenland are unknown. nagssugtoqidian orogen, west greenland the distinction of the nagssugtoqidian orogen in west greenland from the archaean craton to the south was first noted by ramberg (1949). a swarm of basic dykes, the kangâmiut dykes (2040 ma; nutman et al. 1999; mayborn & lescher 2006), which are well preserved in the archaean craton to the south, become increasingly deformed and metamorphosed on entering the nag ssugtoqidian orogen (see fig. 20). this orogen (fig. 2; van gool et al. 2002) extends from søndre strømfjord to disko bugt in west greenland and continues southeastwards beneath the inland ice to the ammassalik region in south-east greenland. it mainly consists of reworked archaean gneisses [74, 75] (connelly & mengel 2000) but also includes palaeoproterozoic supracrustal and intrusive rocks [67, 78, 81]. in west greenland main structures trend ene–wsw, and the orogen exhibits a number of prominent ene-trending shear zones (among which is the nordre strømfjord shear zone; k. sørensen et al. 2006) that separate areas characterised by open folding. the peak of proterozoic tectonic and metamorphic activity was at c. 1850 ma (taylor & kalsbeek 1990) when large parts of the orogen underwent granulite facies metamorphism. high-grade meta2 km pyroxene-amphibolite unit surface of an upper surface of a lower pyroxene-amphibolite unit fig. 9. structural stereogram of the western toqqusap nunaa area north of atammik (c. 65°n), southern west greenland, showing complex fold structures in the late archaean gneiss complex. slightly modified from berthelsen (1960). morphism was followed by an extended period of uplift and cooling (willigers et al. 2002). palaeoproterozoic orogenic activity is believed to be related to collision of two archaean continents, with one or more sutures present within the nagssugtoqidian orogen (kalsbeek et al. 1987; van gool et al. 2002; garde & hollis in press). the reconstruction of st-onge et al. (2009) shown in fig. 3 displays two main sutures, which delimit an interjacent region termed the aasiaat domain. the northwestern part of this region appears only to have been little affected by palaeoproterozoic deformation and metamorphism (e.g. mazur et al. 2006; stendal et al. 2006). tectonically emplaced lenses of metaperidotite of apparent mantle origin are common along thrusts that separate archaean and palaeoproterozoic rocks in the central part of the orogen (kalsbeek & manatschal 1999). supracrustal rocks palaeoproterozoic supracrustal units, dominated by pelitic and semipelitic metasedimentary rocks, are prominent in the central part of the nagssugtoqidian orogen [67]. marble and calc-silicate rocks are common within these units, and pelitic rocks may be rich in graphite. deposition of these units took place between c. 2000 and 1920 ma ago: they are cut by sheets of 1910 ma quartz diorite, and the youngest detrital zircons are c. 2000 ma old (nutman et al. 1999). small islands ne of aasiaat/ egedesminde expose well preserved c. 1850 ma tholeiitic pillow lavas, chloritic and aluminous shales, manganiferous bif etc., interpreted to represent ocean floor and distal turbidite deposits (garde & hollis in press). within the nagssugtoqidian orogen archaean meta sedimentary rocks [68] are also present, for example at the southern shore of disko bugt (hollis et al. 2006). they are not easily distinguished in the field from proterozoic rocks and, since isotopic age determinations were few at the time of map compilation, not all supracrustal sequences shown on the map have been assigned to the correct age category. for example, a metasedimentary unit at ler sletten, south-east of aasiaat/egedesminde (68°24’n, 51°14’w), indicated as archaean on the geological map, has been shown to be 23 fig. 10. nagssugtoqidian orogen. rock face at south side of inner nordre strømfjord/nassuttooq, southern west greenland, showing tectonic contact between pale archaean tonalitic gneisses and overlying dark dioritic gneisses (arfersiorfik quartz diorite) and supracrustal rocks of palaeoproterozoic age. thin slices of marble and calc-silicates occur at the contact and within the archaean gneisses. the height of section is c. 350 m. photo: j.a.m. van gool. 24 palaeoproterozoic in age (c. 1900 ma; thrane & connelly 2006). the involvement of palaeoproterozoic supracrustal rocks in complex fold structures and shear zones in the central part of the belt shows that the deformation was of proterozoic age. felsic and intermediate intrusions only a few granitic and quartz-dioritic intrusive bodies are shown on the map. some are archaean [80], whereas others are of palaeoproterozoic age [78]. a large sheet of quartz diorite [81], the arfersiorfik quartz diorite, dated at 1910 ma, occurs close to the border of the inland ice at 68°n (henderson 1969; kalsbeek et al. 1987; k. sørensen et al. 2006; van gool & marker 2007); it is folded and strongly deformed at its margins, but igneous textures and minerals are preserved in its centre. strongly deformed proterozoic quartz-dioritic to tonalitic rocks (not shown on the map) also occur within reworked archaean gneisses south of the main quartzdiorite body. together they have been interpreted as remnants of a palaeoproterozoic arc, tectonically interleaved with the archaean rocks (fig. 10; kalsbeek et al. 1987; van gool et al. 2002). they range in age from 1920 to 1885 ma (connelly et al. 2000). a large area (c. 30 × 50 km) east and north-east of sisimiut/holsteinsborg is made up of palaeoproterozoic (1910–1870 ma) hypersthene gneisses, the sisimiut intrusive complex (kalsbeek & nutman 1996; connelly et al. 2000; van gool et al. 2002). in the field these cannot easily be distinguished from archaean rocks that occur farther east and, since no age determinations on these rocks were available at the time of map compilation, all the rocks in this area are shown on the map as archaean, overprinted by proterozoic granulite facies metamorphism [75]. nagssugtoqidian orogen in south-east greenland aeromagnetic data show a continuation of the nag ssugtoqidian orogen from west greenland beneath the inland ice to south-east greenland (fig. 11). here palaeo proterozoic orogenic activity has been documented from 64°30´n to 68°n (fig. 2) on the basis of deformation and high-grade metamorphism, up to eclogite facies, of mafic dykes. this region, centred around the town of tassilaq/ammassalik, is dominated by reworked archaean gneisses [74, 75] which were tectonically interleaved with metasedimentary rocks during the palaeoproterozoic (fig. 12; chadwick et al. 1989; kalsbeek 1989). juvenile palaeoproterozoic intrusive rocks [81] and post-tectonic (c. 1680 ma) granite bodies [78] are also present. palaeoproterozoic pelitic metasedimentary rocks [67] are common and locally contain abundant kyanite; thick marble units also occur. archaean anorthositic rocks [85] are present in a few places. a detailed geochronological study of rocks and structures in the ammassalik region has recently been reported by nutman et al. (2008b). before the continuity of the nagssugtoqidian orogenic belt from westto south-east greenland was satisfactorily documented, the latter area was termed the ammassalik mobile belt (kalsbeek 1989), but this term should now be abandoned. palaeoproterozoic plutonic rocks a suite of 1885 ma leuconoritic and charnockitic intrusive rocks [81], the ammassalik intrusive complex (aic), occurs as a row of wnw–ese-trending intrusions at 65°30’n in the centre of the east greenland nag 200 km 65° 70° 70°70° 45°45° 400 200 100 150 50 0 –50 –100 –150 –200 –300 (nt) 45° 70° 40° or og en ketilidian orogen n ag ss ug to qi di an fig. 11. aeromagnetic map of southern greenland showing the continuity of the nagssugtoqidian orogen from west to south-east greenland. data from verhoef et al. (1996). 25 ssugtoqidian orogen (friend & nutman 1989). it was emplaced into a succession of sedimentary rocks, in which it caused widespread anatexis and produced garnet-rich granitic gneisses [71]. the aic is interpreted by nutman et al. (2008b) as a palaeoproterozoic arc, which was caught between archaean crustal units during the nagssugtoqidian collision c. 1870 ma ago. palaeoproterozoic quartz-dioritic to tonalitic intrusions [81] occur locally; one is shown just north of latitude 66°n and has been dated at 1900 ma (nutman et al. 2008b). palaeoproterozoic gneisses, formed locally by deformation of such intrusive rocks, are not distinguished on the map. scattered post-tectonic granite plutons [78] associated with diorite and local gabbro occur over the central part of the east greenland nagssugtoqidian orogen. their isotopic age is about 1680 ma (kalsbeek et al. 1993a), much younger than the peak of tectonic and metamorphic activity in the belt at c. 1870 ma (nutman et al. 2008b). isotopic data show that the granites contain major proportions of crustally derived material (taylor et al. 1984). rinkian fold belt the rinkian fold belt (henderson & pulvertaft 1987; grocott & pulvertaft 1990) lies to the north of the nagssugtoqidian orogen in west greenland between latitudes 69°30´n and 75°n (fig. 2). north of nuussuaq it is characterised by the presence of a several kilometres thick palaeoproterozoic sedimentary succession, the karrat group [62], which overlies and is interfolded with reworked archaean gneisses [74]. it has been difficult in the field to define a distinct boundary between the nagssugtoqidian and rinkian belts. however, fig. 12. nagssugtoqidian orogen in south-east greenland. archaean gneisses reworked during the palaeoproterozoic in granulite and retrograde granulite facies, with supracrustal layers consisting of amphibolite and paragneis. east of northernmost sermilik (c. 66°30´n), southeast greenland. the prominent summit is 1750 m high; relief seen is c. 1500 m. photo: j.c. escher. 26 connelly & thrane (2005) observed a significant change in pb-isotopic compositions of k-feldspar in granitoid rocks across a high-strain belt at c. 69°30´n in disko bugt, and suggested that this belt of strong deformation represents the suture between two large crustal blocks. they proposed that the northern, rinkian block forms part of the rae craton of northern canada, and that the southern, nagssugtoqidian block is the northernmost (de formed) part of the north atlantic craton (fig. 3). the areas around qeqertarsuup tunua/disko bugt and the region north of nuussuaq are described separately below. archaean and palaeoproterozoic supracrustal rocks: disko bugt and nuussuaq the geology of the area around disko bugt has been described in detail by garde & steenfelt (1999) and a geological map at 1:250 000 (garde 1994) is included in their paper. supracrustal rocks occur throughout the area. two representative examples: (1) in north-eastern disko bugt an arcuate archaean greenstone belt consists of basic and acid metavolcanic rocks [68, 66]. the basic rocks, mainly greenschists and meta-pillow lavas, contain a subvolcanic sill complex of gabbros and dole rites [82] (marshall & schønwandt 1999). this belt is intruded by the 2800 ma atâ igneous complex (kalsbeek & skjernaa 1999; see below). (2) another supracrustal belt runs along the south coast of nuussuaq. in contrast to most other supracrustal units this belt was demonstrably deposited upon older gneisses. it consists of mafic and ultramafic metavolcanic rocks with rhyolites, metasedimentary schists, banded iron formation and minor exhalitive rocks with widespread, but low-grade gold min e ralisation. a rhyolite from this belt has yielded an age of c. 2850 ma (connelly et al. 2006). the archaean rocks of north-eastern disko bugt are unconformably overlain by a palaeoproterozoic sedimentary succession [62], the anap nunâ group, which has been correlated with the karrat group farther north (see below). the anap nunâ group consists of a lower unit of marble and orthoquartzite overlain by thick shallow-water siltstones and sandstones. although the rocks are folded, palaeoproterozoic deformation and metamorphism are at a minimum (garde & steenfelt 1999). parts of the succession have suffered intensive metasomatic albitisation (kalsbeek 1992). gneisses and intrusions: disko bugt and nuussuaq reworked archaean gneisses [74] in the disko bugt region are similar to those of the nagssugtoqidian orogen. on nuussuaq they show flat-lying tectonic fabrics and contain anorthosite bodies [85] and dioritic intrusions [82] (garde & steenfelt 1999). north-east of disko bugt, 2800 ma tonalitic rocks of the atâ intrusive complex [80] hardly show any signs of archaean or protero zoic deformation. this complex was emplaced into the arcuate archaean greenstone belt [68] mentioned above and retains many magmatic features (kalsbeek & skjernaa 1999). new geochronological data for the disko bugt region, together with an overview of earlier information, have been presented by connelly et al. (2006). the oldest age (3030 ma) was obtained from a dioritic intrusion within amphibolites on south-eastern nuussuaq. palaeoproterozoic supracrustal rocks north of nuussuaq: the karrat group the geology of the rinkian fold belt in the area north of nuussuaq is depicted on the 1:500 000 geological map of greenland, sheet 4, upernavik isfjord (escher 1985). in this region, the karrat group [62] is widely expo sed over a 400 km coastal stretch north of uummannaq (fig. 2). it was deposited unconformably on archaean crystalline basement rocks between c. 2000 ma (u-pb ages of the youngest detrital zircons; kalsbeek et al. 1998a) and the emplacement of the prøven igneous complex [78] at c. 1870 ma (thrane et al. 2005). the karrat group underwent high-grade metamorphism at relatively low pressure at around 1870 ma (taylor & kalsbeek 1990). the karrat group has been divided into three formations (henderson & pulvertaft 1987). the two lower formations, the mârmorilik formation (up to 1.6 km, garde 1978) and qeqertarssuaq formation (more than 2 km, fig. 13), comprise shelf and rift-type sediments, dominated respectively by marbles and clastic sediments with minor volcanic rocks. these two formations may be correlatives, originally separated by a basement high. at 71°07´n, 51°w the mârmorilik formation hosted the now exhausted black angel lead-zinc mine (thomassen 1991). lead-zinc mineralisation has also been observed at other localities in the region. the upper formation, the nûkavsak formation, with a minimum structural thickness of 5 km, is a typical turbidite flysch succession. extensive tight folding makes estimates of the stratigraphic thickness of the karrat group uncertain. proterozoic sedimentary successions 27 similar to the karrat group occur in the foxe fold belt on the western side of baffin bugt in north-eastern canada (the piling and penhryn groups; henderson & tippet 1980; henderson 1983) suggesting correlation of the rinkian belt of greenland and the foxe fold belt of canada (see fig. 3). connelly et al. (2006) suggest that the karrat group was deposited on the passive margin of the rae craton before collision with the north atlantic craton. the karrat group and its underlying crystalline basement are complexly interfolded into gneiss-cored fold nappes (fig. 14; henderson & pulvertaft 1987) which were subsequently refolded into large dome structures. tectonic interleaving of cover rocks with basement gneisses by thrusting has also taken place so that locally proterozoic supracrustal rocks occur as enclaves within archaean gneisses. the extent of this process is exemplified by an isolated occurrence of pb-zn mineralised marble at 70°30´n, 52°30´w in the centre of nuussuaq (garde & thomassen 1990). gneisses and intrusive rocks north of nuussuaq reworked archaean gneisses [74] north of nuussuaq are similar to those elsewhere. commonly they display flat-lying fabrics related to palaeoproterozoic thrusting. north of nuussuaq sheets of archaean augen gneiss (not distinguished on the map) have been used as structural markers to unravel the complex thrust tectonics of that area (pulvertaft 1986). the 1870 ma prøven igneous complex [78] (thrane et al. 2005) in the upernavik area (c. 72°30´n) consists mainly of charnockitic rocks emplaced into archaean gneisses and metasedimentary rocks of the karrat group, which are here at granulite facies. samples from the prøven igneous complex have an a-type geochemical signature, and isotope data indicate that the magma was formed by anatexis of archaean gneisses and palaeo proterozoic sedimentary rocks at depth. melting is suggested to have been induced by upwelling of hot asthenospheric mantle due to delamination of mantle lithosphere following continental collision in the disko bugt area (see above). north-west greenland and the inglefield orogenic belt the region between 75°15´ and 81°n in north-west greenland is covered by the geological map of greenland, sheet 5, thule (dawes 1991, 2006) and sheet 6, hum boldt gletscher (dawes 2004; dawes & garde 2004). the region up to c. 77°30´n has not been investigated in detail. it consists mainly of reworked archaean gneisses [74] with local amphibolites and banded iron formation. the karrat group has not been recognised in this region. the c. 2700 ma kap york meta-igneous complex [82] at 76°n is composed of a suite of plutonic rocks ranging from gabbro to granite, and a major anorthosite complex [85] is exposed at 77°30´n (nutman 1984). an overview of available geochronological information is given in nutman et al. (2008a). the area between c. 77°30´ and 79°n contains the palaeoproterozoic inglefield orogenic belt which mainly consists of high-grade palaeoproterozoic supracrustal and intrusive rocks that are overlain by mesoproterozoic sedimentary rocks with basaltic sills, the thule supergroup y y y y y y y y y y y y y y y y y y y y y y y y y y y 20 00 m y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y y yy y y y y y y y y y y y karrat isfjord 71°36´n marble dolomite amphibolite pelite semipelite quartzite 71°08´n m âr m or ili k f m q eq er ta rs su aq f m n ûk av sa k f m basement maarmorilik fig. 13. schematic lithostratigraphic sections of the palaeoproterozoic karrat group in the maarmorilik – karrat isfjord region. the nûkavsak formation consists of a flysch succession of interbedded greywacke and mudstone, now metamorphosed at amphibolite facies. based on garde (1978) and henderson & pulvertaft (1987). 28 [3, 4, 5], and by cambrian deposits [23, 25] of the franklinian basin. the inglefield belt is divided into two parts by the e–w-trending sunrise pynt straight belt (spsb, not shown on the map) at c. 78°20´n. archaean rocks south of the spsb have been intruded by c. 1980 ma tonalites and diorites (nutman et al. 2008a). the oldest rocks in inglefield land, north of the spsb, are highgrade metasedimentary rocks of the etah group. most of the group is composed of variably migmatised para gneisses, shown as granulite facies gneisses [71] on the geological map, while better preserved marble-dominated units are shown as supracrustal rocks [67]. zircon geochronology brackets deposition of the etah group between 1980 and 1950 ma (nutman et al. 2008a). the etah group has been intruded by a variety of metaplutonic rocks (not shown on the map), mainly of intermediate to felsic composition, the etah meta-igneous complex. most of these rocks are strongly deformed, but less deformed syenitic and monzonitic rocks are also present, and post-tectonic granites occur locally. dioritic and granitoid rocks were emplaced during several periods, c. 1950–1940 ma and c. 1920 ma, with high-grade metamorphism around 1920 ma, while late granites have ages of 1780 and 1740 ma (nutman et al. 2008a). sm-nd isotopic data show that the older intrusive rocks are of juvenile origin, whereas some of the late intrusions were formed by crustal melting (nutman et al. 2008a). the inglefield belt is interpreted as a palaeo proterozoic orogen, formed by collision of archaean crustal blocks. ketilidian orogen orthogneisses cut by dolerite dykes at the southern margin of the archaean craton are unconformably overlain by palaeoproterozoic sedimentary rocks [64] and basalts [63]. towards the south these supracrustal sequences, together with the underlying archaean gneisses and dykes, are progressively affected by deformation and metamorphism as the ketilidian orogen (fig. 3) is approached. the centre of the ketilidian orogen consists mainly of juvenile palaeoproterozoic granitic rocks, the julianehåb batholith [70, 78]. in the southern part of the orogen high-grade metasedimentary rocks [67] and large intrusions, shown as rapakivi granites [77] on the map, are prominent. the ketilidian orogen is covered by the geological map of greenland, sheet 1, syd grønland (allaart 1975; garde 2007b). during the 1990s the ketilidian orogen was reinvestigated in more detail. a comprehensive report on this new information has been presented by garde et al. (2002). palaeoproterozoic supracrustal rocks in the northern border zone the best preserved ketilidian supracrustal rocks occur in grænseland and midternæs, north-east of ivittuut, where they are locally almost unmetamorphosed and only superficially deformed (fig. 15); the age of deposition is not precisely known. the succession has been divided into (1) a lower sedimentary part, the vallen group, with c. 1200 m of shales and greywackes with grey gneiss dolerite dykescree grey gneiss (allochthonous) karrat group discordant amphibolites dark banded migmatitic gneiss glacier 1692 m e w fig. 14. sketch of part of one of the characteristic palaeoproterozoic nappes in the rinkian fold belt (kigarsima nappe seen from the north). north side of upernivik ø (71°20´n), central west greenland. slightly modified from henderson & pulvertaft (1987). subordinate quartzite, conglomerate and carbonate rocks [64], and (2) an upper volcanic part, the sortis group [63], which consists mainly of basic pillow lavas and contemporaneous basic sills (bondesen 1970; higgins 1970), and has been interpreted to represent ocean floor related to initial rifting. the two groups are in tectonic contact, and it is likely that the sortis group was thrust upon the vallen group. southwards these supracrustal rocks become progressively deformed and intruded by ketilidian granites. palaeoproterozoic granitoids and basic– intermediate intrusions, the julianehåb batholith the central part of the ketilidian orogen is mainly built up of granites, granodiorites and tonalites, commonly with porphyritic textures, collectively known as the julianehåb batholith (‘julianehåb granite’ in older publications). large parts of the batholith were emplaced between 1868 and 1796 ma in a sinistral transpressive setting (chadwick & garde 1996; garde et al. 2002; pulvertaft 2008). major shear zones were formed during emplacement of the batholith, giving rise to tectonic fabrics of variable intensity. the most intensely deformed parts of the batholith are shown as gneisses [70] on the geological map, less deformed varieties as foliated and non-foliated granitic rocks [78]. basic and intermediate intrusions [81] of various ages are also present. these were commonly emplaced simultaneously with felsic magmas, and may occur as mixed rocks in net-veined intrusions. many of the basic and intermediate plutonic rocks are appinites (see fig. 20), i.e. they contain hornblende as the main primary mafic mineral (e.g. pulvertaft 2008). isotopic data show that the julianehåb batholith is of juvenile proterozoic origin (van breemen et al. 1974; patchett & bridgwater 1984; kalsbeek & taylor 1985; garde et al. 2002) and does not represent reworked archaean rocks as previously believed. 29 fig. 15. the basal part of the ketilidian supracrustal succession in central grænseland, south-west greenland, with the archaean basement in the left background, viewed towards west-north-west. the unconformity forms the red-brown slope in the middle distance, with sporadic sub-ketilidian regolith and ketilidian carbonate deposits. the dark layer just above the contact is iron formation. relief is about 500 m. orange tent in the right foreground indicates the scale. photo: a.a. garde. metasedimentary rocks in the south-eastern part of the ketilidian orogen high-grade supracrustal units [67] are prominent in the south-eastern part of the ketilidian orogen. they are composed of psammitic and semipelitic gneisses with local marbles and basic metavolcanic rocks. acid volcanic rocks [65] occur in the inner fjord area north-east of qaqortoq/julianehåb (61°30´n). the clastic sediments are composed mainly of erosion products of the julianehåb batholith, produced more or less contemporaneously with its emplacement; they are interpreted to represent a fore-arc basin. the rocks underwent highgrade, low-pressure metamorphism, up to granulite facies, and widespread anatexis occurred at c. 1790 ma (garde et al. 2002). the ketilidian rapakivi suite flat-lying sheets of rapakivi ‘granite’ [77], folded into kilometre-scale arcs and cusps, are a prominent constituent of the south-easternmost part of the ketilidian orogen (fig. 16). the rocks are characterised by mantled kfeldspar phenocrysts, high fe/mg ratios and high levels of incompatible elements. rather than true granites, the suite mainly includes quartz monzonites, quartz syenites, and norites. isotopic ages between 1720 and 1750 ma have been obtained from these rocks (gulson & krogh 1975; garde et al. 2002). 30 r ms fig 16. part of the ketilidian orogen on the north side of lindenow fjord (c. 60°30´n) viewed towards the east. massive outcrops of a subhorizontal, c. 1740 ma sheet of the rapakivi suite (r) form the steep cliffs to the north-west (left), above flat-lying metasedimentary rocks (ms) [67]. highest summits are about 2500 m. photo: a.a. garde. archaean–palaeoproterozoic basement in the east greenland caledonian orogen the east greenland caledonian orogen is built up of fartravelled allochthonous thrust sheets overlying the eastern margin of the greenland shield. an overview of the geology of the orogen has recently been provided by higgins et al. (2008), and the region is covered by a new geological map at a scale of 1:1 000 000 (henriksen 2003). crystalline basement rocks are prominent both within the thrust sheets and the underlying foreland. they were overlain by neoproterozoic and palaeozoic sedimentary successions prior to involvement in the caledonian orogeny (higgins & leslie 2008). in the scoresby sund region, 70–72°n, archaean basement gneisses with mafic dykes [74] are prominent. north of c. 72°50´n the crystalline basement consists mainly of palaeoproterozoic orthogneisses [70] (kalsbeek et al. 1993b). in the border region, 72–73°n, palaeo proterozoic granitoid rocks have been intruded into archaean gneisses (thrane 2002). an overview over the precambrian evolution of this region is given by kalsbeek et al. (2008a). the archaean basement complex [74] in the inner scoresby sund region and areas immediately to the north, 70°–72°50´n, consists of a variety of migmatitic gneisses (fig. 17) with scattered foliated granitoid plutonic rocks [80]. in the charcot land tectonic window in the northwestern part of the scoresby sund region (72°n) the archaean basement of the foreland is overlain by a palaeoproterozoic supracrustal succession [67] consisting of low-grade metasedimentary and metavolcanic rocks (steck 1971); these are cut by two major postkinematic granodioritic–granitic intrusions [78] emplaced c. 1840 ma ago (hansen et al. 1980). similar rocks [45] occur in the eleonore sø window (c. 74°n), where they are cut by sheets of quartz porphyry, dated at c. 1915 ma (kalsbeek et al. 2008). the basement gneisses in the central and northern parts of the fold belt (north of 74°n), mainly comprise rock units formed c. 2000 ma ago during a palaeo proterozoic event of juvenile crust formation. both older migmatitic gneisses and younger, more homogeneous granites are present. some of the latter have been dated at c. 1750 ma. most of the gneisses are at amphibolite 31 fig. 17. crystalline basement within the caledonian orogen. archaean gneisses with dark amphibolite layers, affected by palaeoand early neoproterozoic orogenic events, and subsequently reworked during the caledonian orogeny. south of inner nordvestfjord/ kangersik kiatteq, scoresby sund region (71°30´n), central east greenland. the profile height is c. 1000 m. 32 facies [70], with occasional areas of granulite facies [71]. large parts of the region underwent caledonian eclogite facies metamorphism (gilotti et al. 2008) which, however, is not registered in the gneisses. supracrustal rocks [67] occur locally. a few isolated intermediate and mafic intrusions [81] occur in the dove bugt region (76–78°n; hull et al. 1994). within this large region of palaeoproterozoic rocks archaean orthogneisses [74] have been documented at two localities: at danmarkshavn (76°40´n; steiger et al. 1976) and in payer land (74°30´n, 23°w; elvevold et al. 2003). the relationships of these archaean rocks with the surrounding proterozoic gneisses are uncertain. archaean–palaeoproterozoic basement beneath the inland ice little is known about the geology of the area now covered by greenland’s central ice sheet – the inland ice (dawes 2009b). however, in 1993 a 1.5 m core of bedrock was retrieved from beneath the highest part of the ice sheet (> 3000 m) at the gisp 2 ice core locality (fig. 2; 72°35´n, 38°27´w). the rock is a leucogranite, and shrimp u-pb zircon data on a few poorly preserved zircons indicate that it is of archaean origin, but strongly disturbed by one or more subsequent tectonometamorphic events, most likely during the palaeoproterozoic (a.p. nutman, personal communication 1995). these results have been confirmed by sm-nd, rb-sr and pbpb isotope data (weis et al. 1997). three samples from ice-transported blocks of granitoid rocks from the area south and south-east of independence fjord, north greenland, have yielded smnd model ages (depaolo 1981) of 3.04–3.38 ga (kalsbeek & frei 2006) and support the view that significant parts of the hidden basement of north-eastern greenland may consist of archaean rocks. 33 the greenland precambrian shield is mainly composed of crystalline gneisses and plutonic rocks older than 1600 ma. younger rock units, mesoproterozoic to phanerozoic in age, are in part related to the formation of sedimentary basins and fold belts along the margins of the stable shield. two major palaeozoic fold belts – the elles merian fold belt of ellesmere island (canada) and north greenland and the caledonian fold belt of east greenland – developed along the north and east margins of the shield respectively. in the descriptions that follow the onshore proterozoic to phanerozoic deposits and orogenic events throughout greenland are presented chronologically within the framework of major depositional basins. palaeoto mesoproterozoic unfolded units independence fjord group, north greenland the earliest recorded major depositional basin developed on the greenland shield is represented by the independence fjord group [31] (figs 18a, b) which is exposed over large areas of eastern north greenland and north-east greenland between north-eastern peary land (83°n) and westernmost dronning louise land (77°n). the group is more than 2 km thick, with its base only exposed in western dronning louise land. the independence fjord group has been studied primarily in the type area around independence fjord in north greenland (see geological map of greenland 1:500 000, sheet 8, peary land; bengaard & henriksen 1986). it is dominated by alluvial clastic deposits, mainly sandstones that form three 300–900 m thick, laterally correlatable units. these are separated by two laterally extensive, much thinner (4–90 m) silt-dominated units that represent deposition in ephemeral lakes. deposition of the independence group took place in an intracratonic sag basin and the development of extensive lacustrine conditions suggests that sedimentation was controlled by basin-wide changes in subsidence rates (collinson et al. 2008). deposition of the independence fjord group took place between the end of the palaeoproterozoic orogenic events in northern greenland at c. 1750 ma and the intrusion of the midsommersø dolerites at 1380 ma (see below). rb-sr dating of clay minerals from siltstones by larsen & graff-petersen (1980) indicated an age for diagenesis at c. 1380 ma, but the coincidence of this age with the time of emplacement of the midsommersø dolerites suggests that this is not the time of sediment deposition. geochronological data on detrital zircons indicate that most of the detritus that formed the inde pendence fjord sandstones was derived from palaeo proterozoic sources (2000–1800 ma; kirkland et al. 2009). the sandstones and siltstones of the independence fjord group are cut by numerous mafic sheets and sills, the ‘midsommersø dolerites’ (kalsbeek & jepsen 1983; kalsbeek & frei 2006), for which a u-pb baddeleyite age of 1382 ± 2 ma has been obtained (upton et al. 2005). the presence of sheets of ‘rheopsammite’ (intrusive rocks formed by partial melting of sandstone at depth; jepsen 1971; kalsbeek & frei 2006) witnesses to the intensity of this magmatic event. although the dolerites form a significant proportion of the outcrop area of the inde pendence fjord group, they are not shown on the present map, but appear on the geological map of greenland 1:500 000, sheet 8, peary land, referred to earlier. they are depicted with other important dyke swarms on fig. 20. zig-zag dal basalt formation, north greenland the mesoproterozoic zig-zag dal basalt formation [30] consists of an up to 1350 m succession of well-preserved tholeiitic flood basalts. its main outcrop area is south of independence fjord in eastern north greenland. the zig-zag dal basalt formation conformably overlies the independence fjord group and is itself disconformably overlain by the hagen fjord group (fig. 24). south of independence fjord the basalt succession crops out over an area of 10 000 km2, but local occurrence of similar basalts in eastern peary land indicates that the formation once covered a large part of north greenland. a close geochemical similarity with the midsommersø proterozoic to phanerozoic geological development after formation of the precambrian shield 34 fig. 18. palaeoproterozoic independence fjord group sandstones. a: undeformed succession on the south side of independence fjord cut by c. 1380 ma midsommersø dolerite intrusions (c. 82°n), eastern north greenland. profile height is c. 800 m. b: folded and metamorphosed sandstones and dolerite sills within the caledonian fold belt (see text). north of ingolf fjord (c. 80°30´n), kronprins christian land, eastern north greenland. profile height is c. 1000 m. b a dolerites implies that the basalts are related to the same igneous event that produced the dolerites, and an age of c. 1380 ma for the basalts is therefore indicated. the zig-zag dal basalt formation is divided into three main units. a ‘basal unit’ of thin aphyric basalt flows is 100–200 m thick and includes pillow lavas in its lower part. the overlying ‘aphyric unit’ (c. 400 m) and the uppermost ‘porphyritic unit’ (up to 750 m) together comprise 30 flows of mainly subaerial lavas. the present distribution pattern of the flows shows a maximum thickness of the succession in the area south of independence fjord, implying subsidence of this central region during the extrusion of the basalts and prior to the peneplanation which preceded deposition of the hagen fjord group. detailed investigations of the basalts have been carried out by kalsbeek & jepsen (1984) and upton et al. (2005). based on trace element and isotope data the latter authors conclude that magma generation took place in an upwelling mantle plume underneath an attenuating continental lithosphere. the lavas of the porphyritic unit are considered to represent essentially uncontaminated plume-source melts. correlation with similar rocks in the northernmost part of the east greenland caledonides sandstones and conglomerates interpreted as strongly deformed representatives of the independence fjord group [31] are found within the northernmost parts of the caledonian fold belt in kronprins christian land and areas to the south (geological map of greenland 1:500 000, sheet 9, lambert land, jepsen 2000; pedersen et al. 2002; collinson et al. 2008). as in the north green land platform, they are cut by numerous sheets of dolerite (fig. 18b). basaltic and andesitic lavas in this area are shown on the map as zig-zag dal basalt formation [30], but shrimp u-pb dating has yielded an age of 1740 ma for associated rhyolitic rocks (kalsbeek et al. 1999), and correlation with the 1380 ma zig-zag dal basalt formation is therefore excluded. the sandstones and conglomerates in kronprins christian land are interbedded with the lavas, and an age of c. 1740 ma is therefore indicated. this age is similar to that of the youngest granites within the crystalline basement in the caledonian fold belt (see p. 31), and the sedimentary rocks can be regarded as molasse-type deposits related to the breakdown of the palaeoproterozoic orogen in north-east greenland. if a correlation with the independence fjord group in the platform is assumed, the sandstones and conglomerates in kronprins christian land must represent the lowermost part of that group. gardar province, south greenland the mesoproterozoic gardar province (upton & emeleus 1987; kalsbeek et al. 1990; upton et al. 2003) is characterised by faulting, deposition of sediments and volcanic rocks, and alkaline igneous activity. an approx imately 3400 m thick succession of sandstones and lavas referred to as the eriksfjord formation (poulsen 1964) accumulated within an ene–wsw-trending continental rift, preserved at about 61°n. within and outside the rift, major central intrusions and numerous dykes were em placed (see also dyke map fig. 20). an intrusive com35 0 500 1000 1500 m w e 2000 m0 1000 augite syenite alkali acid rocks pulaskite and foyaite sodalite foyaite naujaite lujavrite/pegmatitic borders kakortokite/pegmatitic borders ??? fig. 19. diagrammatic cross-section of the ilímaussaq intrusion, gardar province, west of narsaq in south greenland. the intrusion has an outcrop area of 17 × 8 km and has been dated at 1143 ± 21 ma (see review by kalsbeek et al. 1990; h. sørensen 2006a). slightly modified from andersen et al. (1981). 36 dykes in greenland with one trend direction with two or more trends deformed and metamorphosed dolerites and associated dykes legend kimberlites, lamproites and lamprophyric dykes there are few areas in greenland where the rocks are not cut by mafic dykes. the dykes range in age from palaeoarchaean in the godthåbsfjord area to cenozoic in parts of north, east and west greenland. it is very difficult to date mafic dykes, especially where they have been deformed and metamorphosed, and early k-ar and rb-sr age determinations have proved to be imprecise and sometimes entirely misleading. in many cases the age of the dykes is therefore imperfectly known. moreover, in cases where precise age determinations have been carried out, results show that dykes previously believed to belong to a single swarm may have significantly different ages. the diagrams on the opposite page illustrate the history of dyke em placement in greenland, based on the best age estimates available at present. among the best known dyke swarms in greenland are the ameralik dykes in the godthåbsfjord area, which were intruded into eoarchaean gneisses, but are cut by mesoand neoarchaean granitoid rocks; this permits distinction between eoarchaean and mesoand neoarchaean lithologies. the kangâmiut dykes in west green land are well preserved in the archaean craton, but deformed and metamorphosed in the palaeoproterozoic nagssugtoqidian orogen to the north; this makes it possible to monitor the influence of nagssugtoqidian meta morphism and deformation on the host rocks. undeformed proterozoic dolerite dyke belonging to the ‘md’ dyke swarm, cutting archaean orthogneisses, northern fiskefjord region, southern west greenland. photo: a.a. garde. deformed and fragmented archaean ameralik metabasic dyke cutting eoarchaean ‘amîtsoq’ gneisses, godthåbsfjord region, southern west greenland. width of view c. 1.5 m. photo: a.a. garde. fig. 20 (above and facing pages): diagrammatic representation of the major suites of mafic dykes and sills in greenland. compiled by j.c. escher and f. kalsbeek 1997. 37 a dykes and sills related to tertiary rifting, associated with basalts a gardar sills and dykes a dolerite swarm b lamproites c appinites (sensu lato) a kangâmiut dykes b umîvik dyke swarm c ‘md’ dykes a ameralik and tarssartôq dykes (age in ma) c. 15 – c. 60 152–166 c. 65 c. 555–604 c. 135 c. 1100 – c. 1300 c. 670 – c. 1310 c. 1380 c. 1650 c. 1750 c. 1780 – c. 1850 c. 2040 c. 2050 0 ma 500 ma 1000 ma 1500 ma 2000 ma 2500 ma 3500 ma c b sills and dykes b sills and dykes of the nares strait and smith sound groups c. 630 – c. 730 n n n c n n n c. 3460 – c. 3510 (norite) c. 2200 (tholeiite) c. 2150 d b b a c a b a a b b c a a a a a c c b b b a a a b kimberlites and lamprophyres c mafic alkaline dykes a kimberlites and lamprophyres c midsommersø dolerites, associated with basalts d ‘td’ dykes plex at the head of danell fjord (c. 60°50´n, 43°30´w) in south-east greenland is indicated on the map as ketilidian rapakivi granite [77], but later radiometric dating has yielded a gardar age (garde et al. 2002). the sedimentary [12] and volcanic rocks [11] of the eriksfjord formation rest uncomformably on ketilidian granites. the eriksfjord formation comprises c. 1800 m of sedimentary strata and 1600 m of volcanic rocks. the sedimentary rocks, mainly found in the lower part of the succession, are fluvial and aeolian arkosic to quartzitic sandstones and conglomerates (clemmensen 1988; tirsgaard & øxnevad 1998). the volcanic rocks are dominated by basaltic lavas, with subordinate trachytes and phonolites in the upper part and a carbonatite complex in the lower part (stewart 1970; larsen 1977; upton & emeleus 1987). the age of the eriksfjord formation is c. 1170–1200 ma (paslick et al. 1993). the gardar intrusive complexes [56] range in age from c. 1300 to c. 1120 ma and have been divided into three age groups (upton & emeleus 1987; upton et al. 2003). they comprise central ring intrusions, complexes with several individual intrusive centres, and giant dykes (emeleus & upton 1976; upton & emeleus 1987). petrologically, the intrusive complexes are dominated by differentiated salic rocks including syenites, nepheline syenites, quartz syenites, and granites (fig. 19); mildly alkaline gabbros and syenogabbros are subordinate but are dominant in the giant dykes. the intrusions were em placed in the middle part of the gardar rift as well as in the areas to the north-west and south-east. major swarms of basic dykes of gardar age occur throughout south and south-west greenland (see dyke map, fig. 20). early neoproterozic orogenic units reworked in the east greenland caledonian fold belt a suite of early neoproterozoic augen granites and leu cogranites [55] is widely distributed within the krum medal supracustal sequence of the high-grade uppermost caledonian (hagar bjerg) thrust sheet between scoresby sund (70°n) and about 74°n; the granitoids have yielded protolith ages of 940–910 ma (jepsen & kalsbeek 1998; kalsbeek et al. 2000; watt & thrane 2001). these magmatic events are contemporaneous with high-grade metamorphism dated in overgrowth rims on detrital zircons (kalsbeek et al. 2000; watt et al. 2000; watt & thrane 2001), as well as ductile deformation that, at least locally, produced nappe-scale recumbent folds in the reworked migmatite and paragneiss complex [52]. a comparable scenario is recorded in eastern svalbard where 970–940 ma events are recorded and augen granites have been emplaced synchronously with deformation (johansson et al. 2000); in scotland zircon geo chronology has revealed a range of tectonothermal events from 840–730 ma (leslie et al. 2008). supracrustal rocks the krummedal supracrustal sequence [46] consists of a 2500–8000 m thick suite of pelitic, semipelitic and quartzitic rocks generally metamorphosed at amphibolite facies (henriksen & higgins 1969; higgins 1974, 1988; higgins & leslie 2008; figs 21, 22). lateral and vertical lithological variations are considerable and correlation between the various local successions has not been possible. contacts with the underlying archaean [74] and palaeoproterozoic gneisses [70] are generally 38 + + + + + + + + + + + 20 00 m quartzite semipelite pelite augen granite migmatitic thrust contact fig. 21. sections of the mesoproterozoic krummedal supracrustal sequence, north of inner nordvestfjord/kangersik kiatteq (71°30´n), scoresby sund region, central east greenland. based on higgins (1974). 39 conformable, but rare discordances may reflect preservation of an original unconformity (higgins et al. 1981). the ‘smallefjord sequence’ [46] that crops out between grandjean fjord (75°n) and bessel fjord (76°n) (friderichsen et al. 1994) is comparable in lithology and development to the krummedal succession. age determinations on zircons suggest deposition of both sequences later than c. 1100 ma, and high-grade metamorphism dur ing an early neoproterozoic event at c. 950 ma (strachan et al. 1995; kalsbeek et al. 1998b). the krummedal sequence of the lowermost caledonian (niggli spids) thrust sheet appears to lack the early neoproterozoic granitoids and migmatitic developments recorded in similar rocks within the uppermost (hagar bjerg) thrust sheet (see later). migmatites and granites the krummedal supracrustal sequence of the uppermost hagar bjerg thrust sheet in the southern part of the east greenland caledonian fold belt has been intensely migmatised and transformed into paragneiss [52], and as noted above, has been intruded by sheets of augen granites up to 1000 m thick as well as other granite bodies [55] (steiger et al. 1979). in the scoresby sund region these rock units have been deformed into major recumbent folds (leslie & nutman 2003). a second generation of caledonian granite intrusions [54] was produced by partial melting of the krummedal supracrustal sequence, and some of these granites migrated upwards into the overlying eleonore bay supergroup [44]. mesoproterozoic – early neoproterozoic sedimentary basin in north-west greenland and ellesmere island thule supergroup the thule basin is defined by a thick undeformed sedimentary-volcanic succession – the thule supergroup – that straddles northern baffin bugt and smith sund (dawes et al. 1982; dawes 1997, 2004, 2006). the eastern and western parts of the basin are exposed in northwest greenland and south-eastern ellesmere island (canada), with extensive sections offshore (funck et al. 2006). as such, the thule basin is one of several meso– neoproterozoic intracratonic depocentres fringing the northern margin of the canadian–greenland shield. in greenland, the rocks are widely exposed between ingle field land (79°n) and thule air base/pituffik (76°n). the thule supergroup has a cumulative thickness of at least 6 km and comprises continental to shallow marine sedimentary rocks, basaltic rocks and a conspicuous numfig. 22. krummedal supracrustal sequence comprising rusty garnet iferous gneissic schists and siliceous paragneisses, inner nordvestfjord/ kangersik kiatteq (71°30´n), scoresby sund region, central east greenland. profile height is c. 1500 m. ber of doleritic sills. resting with a profound unconformity on the peneplained archaean–palaeoproterozoic crystalline shield, the basin developed between c. 1270 ma and around 900 ma ago (for discussion, see dawes 1997, 2006; samuelsson et al. 1999). it is dissected by a half-graben system dominated by wnw–ese-trending faults. the thule supergroup is divided into a lower part of three groups [5] and an upper part of two [3, 4]. all of the groups contain red beds. when the map was compiled, a middle – late proterozoic age was assigned (dawes & vidal 1985; dawes & rex 1986), but reappraisal of the acritarch fauna suggests a middle mesoproterozoic to early neoproterozoic age for the entire succession (samuelsson et al. 1999; dawes 2006). the lower part comprises: (1) the smith sound group of mainly shallow marine sandstones and multicoloured shales with stromatolitic carbonates; (2) the nares strait group which at its base consists of inner shelf mudstones and fluvial sandstones, succeeded by terrestrial basaltic extrusive rocks and volcaniclastic red beds overlain by stromatolitic carbonate and shales topped by shallow marine sandstones; (3) the baffin bay group of multicoloured sandstones and conglomerates with intervals of shale–siltstone, mainly of mixed continental to shoreline origin. the upper part of the thule supergroup comprises: (4) the dundas group [4] of deltaic to coastal plain deposits, dominated by dark shales, siltstones and finegrained sandstones with thin carbonate-rich beds, and (5) the narssârssuk group [3], representing deposition in a low-energy environment, with a cyclic carbonate and red-bed siliciclastic succession. the latter comprises interbedded dolomite, limestone, sandstone, siltstone and shale with evaporites. the narssârssuk group, the youngest unit, has a very restricted occurrence in a half-graben on the south-eastern margin of the thule basin (fig. 23). neoproterozoic sedimentary basins in north, north-east and east greenland hagen fjord group, north greenland neoproterozoic basin deposits laid down between 800 and 590 ma ago occur extensively in eastern north green land, where they crop out over an area of 10 000 km2 west of danmark fjord. these deposits, assigned to the hagen fjord group (figs 24, 25), overlie sandstones of the palaeoproterozoic independence fjord group and basalts of the mesoproterozoic zig-zag dal basalt formation together with their (1380 ma) correlatives the midsommersø dolerite intrusions (sønderholm & jepsen 1991; clemmensen & jepsen 1992; sønderholm et al. 2008). the easternmost occurrences of the succession are represented in the caledonian vandredalen thrust sheet in kronprins christian land that has a demonstrable westward displacement of 35–40 km (higgins et al. 2001a, b, 2004b; leslie & higgins 2008). 40 kap york q kt/rf ch jh/bb cc n w rb sb ka ph cc inglefield land kap alexander 50 km ne nw se narssârssuk group smith sound group dundas group precambrian shield baffin bay group nares strait group q: rf: robertson fjord fm kt: kap trautwine fm w: nares strait group ch: clarence head fm bb: jh: cc: cape combermere fm n: smith sound group sb: sonntag bugt fm rb: rensselaer bay fm ka: kap alexander fm ph: pandora havn fm cc: cape camperdown fm baffin bay group qaanaaq fm wolstenholme fm northumberland fm josephine headland fm barden bugt fm fig. 23. cross-section through the thule basin, north-west greenland, with the lower thule supergroup as basin fill, showing the relationships of groups and their formations. the spatial relationship of the dundas and narssârssuk groups superimposed on this mesoproterozoic evolutionary stage is shown by the green bars. vertical exaggeration × 25. slightly modified from dawes (1997). the hagen fjord group [27] has a maximum thickness of 1000–1100 m and comprises a succession of siliciclastic and carbonate sedimentary rocks deposited on a shallow-water shelf. its lower part mainly comprises sandstones which are overlain by a sandstone-siltstone association. the upper part is characterised by limestones and dolomites with abundant stromatolites (fyns sø formation). the overlying sandstone unit (kap holbæk formation) is now known to be early cambrian, and is excluded from the hagen fjord group (smith et al. 2004). the age of the hagen fjord group is poorly constrained, but a pre-600 ma age is now suggested (sønderholm et al. 2008). the rivieradal sandstones [29], now the rivieradal group (smith et al. 2004), are confined to the allochthonous, caledonian, vandredalen thrust sheet in the kron prins christian land area, and are interpreted as deepmarine deposits equivalent in age to the lower part of the hagen fjord group (clemmensen & jepsen 1992; sønderholm et al. 2008). this succession is 7500 to 10 000 m thick and comprises conglomerates, sandstones, turbiditic sandstones and mudstones that accumulated in a major east-facing half-graben basin; the bounding western fault was reactivated as a thrust during the caledonian orogeny (higgins et al. 2001b). a succession of diamictites and sandstones up to 200 m thick, believed to be late precambrian (marinoan, c. 635 ma) in age, forms isolated small outcrops in eastern north greenland; these are known as the morænesø formation [28]. the formation is not included in the redefined hagen fjord group of clemmensen & jepsen (1992) but is in part equivalent or slightly younger in age (collinson et al. 1989; sønderholm & jepsen 1991; smith & rasmussen 2008). eleonore bay supergroup, east and north-east greenland the eleonore bay supergroup comprises a more than 14 km thick succession of shallow-water sedimentary rocks which accumulated in a major sedimentary basin exposed between latitudes 71°40´ and 76°n in east and north-east greenland (sønderholm & tirsgaard 1993; sønderholm et al. 2008). exposures only occur within the present caledonian fold belt, and in general the sedimentary rocks are moderately deformed and weakly to moderately metamorphosed. the nature of the lower contact of the eleonore bay supergroup has been widely debated. the oldest sedimentary rocks are in contact with the krummedal supracrustal succession, with the contact described in some areas as an extensional detachment (hartz & andresen 1995; andresen et al. 1998; white et al. 2002), and in other areas as a westward directed thrust (higgins & leslie 2008; leslie & higgins 2008). relationships are complicated by extensive anatexis and the presence of caledonian granites. sedimen tation is constrained to the interval between c. 900 ma and c. 665 ma by the youngest ages on detrital zircons from the lowest levels of the eleonore bay supergroup and the marinoan (c. 635 ma) age of the overlying tillite group (sønderholm et al. 2008). 41 100 km nw f ra nk lin ia n b as in c am br ia n se kronprins christian land wandel valley fm tavsens iskappe gp brønlund fjord gp buen fm portfjeld fm independence fjord gp (palaeoproterozoic) morænesø fm zig-zag dal basalt fm hagen fjord gp rivieradal group (allochthonous) > 1 00 0 m n eo pr ot er oz oi c o rd ov ic ia n p al ae o– m es opr ot er oz oi c basaltic flowssandstone with intrusives sandstone siltstone kh wandel dal carbonate fig. 24. schematic cross-section of the proterozoic–ordovician succession in eastern north greenland between wandel dal (c. 82°n) and kronprins christian land (c. 80°n). the cross-section shows the relationships between the mesoproterozoic zig-zag dal basalt formation, the neoproterozoic hagen fjord group with correlatives, and the underlying and overlying sequences. colours correspond to those used on the map. bold lines represent erosional unconformities. slightly modified from clemmensen & jepsen (1992). kh, kap holbæk formation. 42 the lower part of the eleonore bay supergroup (fig. 26) consists of up to 9000 m of sandstones, siltstones and minor carbonates assigned to the nathorst land group [44]; these were deposited in a shelf environment with facies associations indicating outer to inner shelf environments (smith & robertson 1999; sønder holm et al. 2008). the upper part comprises three groups (lyell land, ymer ø and andrée land groups) depicted on the map by a single colour division [43]. alternating sandstones and silty mudstones of the lyell land group (fig. 27) reflect deposition in marine shelf environments (tirsgaard & sønderholm 1997; sønderholm et al. 2008). individual units are 40–600 m thick with a total thickness of 2800 m. the overlying 1100 m thick ymer ø group records two significant phases of shelf progradation. depositional environments range from siliciclastic basinal and slope deposits through carbonate slope and shelf deposits to inner shelf siliciclastics and evaporites (sønderholm & tirsgaard 1993; sønderholm et al. 2008). the latest stage of basin fill is mainly represented by the up to 1200 m thick andrée land group of bedded limestone and dolomites, with 10–30 m thick units of stromatolitic dolomite. deposition took place in a carbonate ramp system, with a steepened ramp towards the deep sea to the north-east and a sheltered inner lagoon behind an inner shallow-barrier shoal (frederiksen & craig 1998). the uppermost sequence heralding the marinoan glaciation of the tillite group consists of a strongly retrogradational succession indicating drowning of the carbonate platform and deep marine deposition, followed by a short period of carbonate platform progradation (sønderholm et al. 2008). tillite group, east greenland the tillite group [42] consists of a 700–800 m thick suc cession of marinoan–ediacaran age (c. 635 – c. 575 ma) and includes two marinoan glacigene diamictite formations (hambrey & spencer 1987; sønderholm et al. 2008). it crops out in east greenland between latitudes 71°40´ and 74°n where it overlies the eleonore bay supergroup with no major hiatus, but locally with an erosional unconformity. the tillite group is subdivided into five formations which include sandstones, shales and dolostones in addition to the diamictite formations. kb kb fs cd ja pf fig. 25. hagen fjord group on the north-west side of hagen fjord, eastern north greenland. a lower light coloured sandstone (jyske ås formation, ja, 400 m) is overlain by a multicoloured sandstone-siltstone association (campanuladal and kap bernhard formations, cd and kb, 450 m), with a light coloured limestone–dolomite succession at the top (fyns sø formation, fs, 170 m, and portfjeld formation, pf). the fault has a displacement of c. 300 m down to the right (north). the section is c. 600 m high. from sønderholm et al. 2008. 43 isolated occurrences of diamictites correlated with the tillite group directly overlie crystalline basement complexes in gåseland (70°15´n), charcot land (71°52´n) and in the målebjerg window (73°38´n), all located in the caledonian foreland (henriksen 1986; moncrieff 1989; smith & robertson 1999); the first two of these are shown on the map by a special symbol [41]. these foreland tillites are directly overlain by cambrian quartzites or truncated by the caledonian sole thrust. sedimentary rocks of unknown age in the east greenland caledonides two successions of low-grade metamorphic rocks occur in the nunatak region between 70° and 74°n underlying caledonian thrusts. their correlation with other known successions was uncertain when the map was compiled, and they have been indicated on the map as of ‘unknown age’ [45]. one succession crops out in the gåseland window in the south-west corner of the scoresby sund region (70°15´n), overlying archaean crystalline basement rocks. a thin sequence of weakly metamorphosed marbles and chloritic schists, often highly sheared adjacent to the caledonian sole thrust, overlies diamictites [28] preserved in erosional depressions in the gneiss surface (phillips & friderichsen 1981). the diamictites are now correlated with the marinoan tillite group of the fjord zone (moncrieff 1989), suggesting the overlying sheared marbles and schists are either of early palaeozoic age or belong to a thin lowermost thrust assemblage of diverse lithologies distinguished as the gemmadal thrust sheet in the central part of the fjord region (c. 73°30´n; higgins & leslie 2008). the second succession, traditionally known as the ‘eleonore sø series’, crops out in arnold escher land (74°n; katz 1952). field studies in 1997 have shown the succession to occur in a tectonic window beneath cale donian thrust units of metasedimentary rocks and gneisses. the eleonore sø series comprises low-grade metamorphic sandstones, shales and carbonates associated with volcanic rocks (tuffs and pillow lavas). u-pb ion probe studies on zircons from a quartz porphyry intruding the eleonore sø series indicate a minimum emplacement age of 1915 ± 16 ma (kalsbeek et al. 2008). this succession is overlain unconformably by a thick cambrian quartzite unit which preserves abundant skolithos (slottet formation) and a few hundred metres of lower palaeozoic carbonates (målebjerg formation), recently described and defined by smith et al. (2004). nathorst land group fault / shear zone grenvillian? basement ymer ø group lyell land group andrée land group tillite group limestone dolostone mudstone sandstone 0 2 4 e le o n o r e b a y s u p e r g r o u p 6 8 10 12 14 km fig. 26. schematic composite section of the neoproterozoic eleonore bay supergroup, central fjord zone (72–74°n), north-east greenland. units on the map: nathorst land group [44]; lyell land, ymer ø and andrée land groups [43]; tillite group [42]. based on sønderholm & tirsgaard (1993) and sønderholm et al. (2008). 44 carbonatites, kimberlites and associated rocks, west greenland in addition to the 3007 ma tupertalik carbonatite [84] mentioned earlier, two younger occurrences of carbonatite are shown on the geological map, the 565 ma sarfartoq carbonatite complex [61] south of søndre strømfjord at 66°30´n (secher & larsen 1980), and the c. 165 ma (middle jurassic) qaqarssuk carbonatite complex [59], east of maniitsoq/sukkertoppen at 65°23´n (knudsen 1991). since the compilation of the map another occurrence of carbonatitic rocks has been detected within the archaean craton, the 158 ma tikiusaaq carbonatite complex at 64°n, 49°46´w, c. 100 km east of nuuk/godthåb (steenfelt et al. 2006). carbonatites also occur within some of the intrusive complexes in the mesoproterozoic gardar province (upton et al. 2003). a review of all alkaline-ultramafic and carbonatitic rocks in west greenland (except the newly discovered tikuusaaq occurrence) has been presented by larsen & rex (1992). these rocks are invariably related to episodes of continental rifting. they were formed from small melt fractions generated deep within the lithospheric mantle, and many dykes contain xenoliths of both mantle and crustal origin. most carbonatite complexes are associated with swarms of ultramafic dykes, kimberlites, aillikites etc. (nielsen et al. 2009), here collectively termed ultramafic lamprophyres (sensu lato) (uml dykes). the dykes are too small to be shown on the geological map, but they are important for diamond exploration. hundreds of diamonds have been recovered from a single dyke in the sarfartoq region, the largest of which are c. 4 carats and of good gem quality (hutchison & heaman 2008). a regional overview of diamond occurrences in southern west greenland is given by jensen et al. (2004). recent investigations of uml dykes have concentrated on the areas around sisimiut/holsteinsborg, sarfartoq and maniitsoq/sukkertoppen, the ‘diamond province’ of southern west greenland (nielsen et al. 2009). they fall in several age groups (secher et al. 2009). uml dykes in the sisimiut region have yielded ages of c. 590 ma (scott 1981). dykes in a wide region around the sarfartoq carbonatite complex have neoproterozoic ages between 604 and 555 ma. similar ages were found for dykes in the maniitsoq region, but samples collected around the qaqarssuk carbonatite complex are of jurassic age 152–166 ma. only neoproterozoic dykes have as yet proved to be diamondiferous (secher et al. 2009). fig. 27. part of the upper eleonore bay supergroup, west side of ymer ø (c. 73°n), north-east greenland. succession is approximately 2 km in thickness and includes from left to right: lyell land group (apart from the two lowest formations) and to the right of the black dashed line ymer ø group (the lowest five of seven formations). photo: m. sønderholm. the palaeozoic franklinian basin of north greenland and ellesmere island the palaeozoic franklinian basin extends from the canadian arctic islands across north greenland to kronprins christian land in eastern north greenland, an e–w distance of 2000 km (peel & sønderholm 1991); only part of the canadian segment of the basin is represented on the map. the preserved part of the succession shows that deposition in this e–w-trending basin began in the latest precambrian and continued until at least the earliest devonian in greenland and later devonian to earliest carboniferous in canada; sedimen tation was brought to a close by the midto late palaeozoic ellesmerian orogeny. in the canadian arctic islands deposition continued more or less continuously throughout the devonian and probably into the earliest car boniferous. deposition of clastic sediments of middle and late devonian age in the southern part of the franklinian basin in the canadian arctic islands reflects an early orogenic event with uplift and erosion starting in latest silurian time (trettin 1991, 1998). deposition in the franklinian basin in north green land took place along a passive continental margin, and its evolution during the early palaeozoic resulted in a distinctive differentiation into a southern, broad, shallow shelf bordered to the north by a slope with moderate water depths and a broad deep-water trough (higgins et al. 1991). the shelf succession is dominated by carbonates and reaches 3 km in thickness, whereas the trough deposits are dominated by siliciclastic rocks and have a total thickness of c. 8 km (fig. 28). the shelf–trough boundary was probably controlled by deep-seated faults, and with time the trough expanded southwards to new fault lines, with final foundering of the shelf areas in the silurian. the sedimentary successions in the north green land and canadian (ellesmere island) segments of the basin show close parallels in development, although different lithostratigraphic terminologies are employed (trettin 1991, 1998). the evolution of the franklinian basin in north green land has been divided into seven stages, with significant changes in the sedimentary regime linked to southward expansion of the basin margin (higgins et al. 1991; henriksen & higgins 2000). uppermost neoproterozoic – silurian in north greenland the oldest shelf deposits range from latest neoproterozoic to cambrian in age, and consist of a mixture of carbonates and siliciclastic sediments [25]; they crop out in a narrow, almost continuous zone extending from dan mark fjord in the east through southern peary land to southern wulff land in the west (ineson & peel 1997). the southernmost outcrops farther to the west in ingle field land rest on crystalline basement. three principal divisions are recognised: a lower varied sequence of sandstones, dolomites and mudstones (skagen group), a middle dolomitic unit locally with stromatolites (portfjeld formation), and an upper siliciclastic unit (buen for mation). total thickness reaches 1–2 km. the buen formation in north greenland is noted at one location for its well-preserved soft-bodied fossil fauna (conway morris & peel 1990, 1995, 2008). early cambrian deep-water turbidite trough sediments [26] dominate the northernmost parts of greenland bordering the arctic ocean, and they also crop out in a broad e–w-trending belt north of lake hazen in ellesmere island. the lower part (nesmith beds in canada, paradisfjeld group in greenland) comprises calcareous mudstones and dolomites with, in greenland, carbonate conglomerates at the top. the upper division (polkorridoren group) is made up of thick units of sandy turbidites and mudstones. the thickness of these two divisions totals about 3–4 km (friderichsen et al. 1982; higgins et al. 1991). carbonate sedimentation resumed on the platform in the late early cambrian (ineson et al. 1994; ineson & peel 1997) and continued with minor siliciclastic intervals until the early silurian, giving rise to an up to 1500 m thick succession of carbonate lithologies (brøn lund fjord, tavsen iskappe, ryder gletscher and morris bugt groups, and petermann halvø and ymers gletscher formations [23]). throughout the period sedimentation was influenced by differential subsidence and southwards expansion of the deep-water trough. uplift in eastern north greenland led to erosion of the cambrian to late early ordovician succession in kronprins christian land, after which the middle ordovician to early silurian platform succession was deposited. a broad zone of outcrop can be traced from danmark fjord to washington land, with outliers to the south-west in northern inglefield land. on ellesmere island extensive outcrops are found on judge daly promontory. the up to 1500 m thick succession (fig. 29) of massive dolomites, carbonate grainstones, carbonate mass-flow deposits and evaporites reflects both progradation and aggradation phases of platform evolution. the cambrian – early silurian starved slope and trough deposits (surlyk & hurst 1984) are represented by a condensed succession, dominated for the most part 45 by carbonate mudstones and carbonate conglomerates in the lower part (vølvedal group) and by cherts and cherty shales in the upper part (amundsen land group) [24]. in central north greenland thin-bedded turbidites characterise both the lower and upper parts of the succession. in greenland these deposits occur in thrust slices and anticlinal fold cores (fig. 30; soper & higgins 1987, 1990); in ellesmere island they occur mainly in scattered anticlinal fold cores. thicknesses vary greatly, from a minimum of 50–150 m to a maximum of about 1 km. 46 turbiditic siltstones and sandstones open marine platform carbonates restricted platform carbonates biostromal carbonates platform margin carbonates backmound and intermound carbonates outer shelf-slope carbonates evaporites mounds inner shelf sandstones and shales outer shelf shales and sandstones hemipelagic mudstones and siltstones unconformity f f f f f f f f f f f f f f f f f f f f f f f ff f f f f f f f f cherty shales conglomerates mega breccia proterozoic sandstones and shales crystalline basement rocks terrigenous mudstones f f in gl ef ie ld l an d w as hi ng to n l an d h al l l an d w ul ff la nd n ar es l an d w es te rn p ea ry l an d ce nt ra l p ea ry l an d no rt he rn p ea ry l an d 7t 6t 5t 4t 3t 2t 1 k ro np rin s c hr is tia n la nd 7t 6t 5s 4s 3s 2 1 6s 5s 4s 4s 3s 2s fig. 28. block diagram illustrating relationships between shelf, slope and trough sequences in the lower palaeozoic franklinian basin of north greenland. the schematic fence diagram covers a region of c. 700 km east–west and c. 200 km north–south. shelf stages (s) and trough stages (t) are divided into time intervals. 1: late neoproterozoic? – early cambrian; 2: early cambrian; 3: early cambrian; 4: late early cambrian – middle ordovician; 5: middle ordovician – early silurian; 6: early silurian; 7: later silurian. units on the map: portfjeld and buen formations [25] – stages 2–3 s; brønlund fjord, tavsen iskappe, ryder gletscher and morris bugt groups and petermann halvø and ymers gletscher formations [23] – stages 4–5 s; washington land group exclusive above mentioned formations [21] – stages 6–7 s; skagen, paradisfjeld and polkorridoren groups [26] – stages 1, 2–3 t; vølvedal and amundsen land groups [24] – stages 4–5 t; peary land group [22] – stages 6–7 t. modified from higgins et al. (1991) and with information from m. sønderholm (personal communication 1998). silurian carbonate ramp and rimmed shelf deposits (washington land group [21]) crop out in an almost continuous narrow strip extending from kronprins christian land in the east to washington land in the west (hurst 1980, 1984; sønderholm & harland 1989; higgins et al. 1991). the comparable deposits of this age in ellesmere island have been included in an extension of unit [23] – see legend. sedimentation on the platform was closely linked to the dramatic increase in deposition rates in the trough and was initiated in the early silurian (early late llandovery) by a major system of sandstone turbidites (peary land group [22]) derived from the rising caledonian mountains in the east (hurst & surlyk 1982; surlyk & hurst 1984; larsen & escher 1985). loading effects of the turbidites led to downflexing of the outer platform and expansion of the trough. with progressive drowning of the shelf, carbonate deposition was only locally maintained on isolated reef mounds up to 300 m high (e.g. samuelsen høj formation, hauge bjerge formation). mound formation terminated over much of the region during the late llandovery, but persisted in western north greenland into the late silurian (early ludlow). the silurian turbidite trough deposits occur in a broad belt traceable across north greenland (peary land group [22]) and ellesmere island; the commencement of turbidite sedimentation was essentially synchronous in both regions within the limits of biostratigraphic resolution (trettin 1991). the trough sediments represent the deposits of a major e–w-trending sand-rich turbidite system. palaeocurrent directions in north greenland indicate a source area in the rising mountains of the cale donian fold belt to the east (higgins et al. 1991; henriksen & higgins 2000), whereas current directions in ellesmere island demonstrate an additional source area in the north. the initial phase of sandstone turbidite deposition in north greenland laid down between 500 and 2800 m of sediment within the early silurian (late llandovery); this filled the deep-water trough, buried the former shelf escarpment and led to deposition of black mudstone over extensive former shelf areas. renewed prograding fan systems built up and turbidite deposition continued throughout the silurian, punctuated by an episode of chert conglomerate deposition in the middle wenlockian (surlyk 1995). palaeontological evidence from the youngest deposits in north greenland indicates a late 47 fig. 29. cambro-ordovician platform margin sequence in the foreground and ordovician shelf sequence clastic rocks in the middle distance. view from the south, inner j.p. koch fjord, central north greenland. profile height in the foreground is c. 500 m. f silurian (pridoli) to early devonian age (bendix-almgreen & peel 1974; blom 1999). in ellesmere island this phase of turbidite deposition persisted into the lower devonian; farther to the west in the canadian arctic islands clastic sedimentation associated with the advance of ellesmerian deformation continued through the devonian into the earliest carboniferous. proterozoic–silurian exotic terrane of ellesmere island (pearya) the geological province of pearya, now recognised as an exotic terrane, is confined to northernmost ellesmere island (trettin 1991, 1998). pearya appears to have been accreted during the latest ordovician to early silurian, 48 4 2 0 2 4 6 n km kap washington group volcanics silurian turbiditic sandstones and mudstones, peary land group ordovician fine-grained clastic sedimentary rocks, vølvedal and amundsen land groups cambrian sandstones and mudstones, polkorridoren group cambrian lime mudstones, paradisfjeld groupcretaceous sedimentary rocks wandel sea basin deep-water basin succession cambrian quartzitic sandstones, skagen groupcarboniferous sedimentary rocks fig. 30. middle ordovician – lower silurian sediments in the deep-water sequence of the franklinian basin (amundsen land group dark unit; merqujôq formation light coloured unit). the sedimentary rocks were folded into south-facing tight folds during the ellesmerian orogeny. northeast cape of victoria fjord, central north greenland, view towards the east. profile height is c. 400 m. fig. 31. n–s structural cross-section through the north greenland (ellesmerian) fold belt and its southern foreland in north greenland (c. 39°w, westernmost peary land). compiled from soper & higgins (1990) and henriksen (1992). and underwent further convergence or accretion during the late silurian (de freitas et al. 1999; tessensohn & roland 2000). on the geological map it is represented by two divisions: late mesoproterozoic crystalline rocks [52] and a neoproterozoic to late silurian complex of undifferentiated metasedimentary and metavolcanic rocks (the mainly exotic terrane [2] of the map legend). the rocks of the latter division are stratigraphically or structurally associated with the formation of the franklinian deep-water basin. the crystalline rocks consist of granitoid gneisses and lesser amounts of amphibolite, schist, marble and quartzite in several outcrop areas with different structural settings and trends. the younger su pra crustal complexes include different carbonate and clastic sediments together with varied acid and mafic volcanic rocks. these supracrustal rocks have been folded and constitute the markham fold belt, which is a complex region that fringes the pearya terrane on the southeast. the pearya exotic terrane is noted for emplacement of granite plutons associated with the early middle ordovician m’clintock orogeny, not recorded elsewhere in ellesmere island. ellesmerian orogeny in north greenland and ellesmere island the palaeozoic ellesmerian orogeny, which brought sedimentation in the franklinian basin to a close, involved compression of the lower palaeozoic trough succession against the carbonate shelf to the south following collision with an unknown continent to the north (higgins et al. 2000). the resulting ellesmerian fold belts of both north greenland and northern ellesmere island are characterised by e–wto ne–sw-trending chains of folds, broadly parallel to the main facies boundaries within the franklinian basin. in the north greenland fold belt deformation is most intense in the north, where three phases of folding are recognised and metamorphic grade reaches low amphibolite facies. deformation decreases southwards, and the southern part of the fold belt is a thin-skinned fold and thrust zone (soper & higgins 1987, 1990; higgins et al. 1991) that coincides with the region which was transitional between the platform and trough for much of the cambrian (fig. 31). a prominent belt of major folds is traceable between northern nyeboe land and j.p. koch fjord, and farther east spectacular imbricate thrusts occur north of the head of frederick e. hyde fjord (pedersen 1986). the same general pattern of ellesmerian deformation is seen in ellesmere island, except that the southernmost belt of folding propagated some 100 km southward into the platform, producing the large-scale, concentric-style folding seen north-west of kennedy kanal. the age of the ellesmerian orogenic deformation in north greenland is not well constrained, but is assumed to be late devonian to early carboniferous. 49 0 5 10 km s shelf_slope succession silurian carbonates cambro-ordovician shelf carbonates cambrian sandstones and mudstones, buen formation cambrian dolomites, portfjeld formation precambrian basement fault thrust 50 lower palaeozoic of east greenland cambrian–ordovician sediments in the caledonian fold belt cambrian–ordovician rocks [40] make up an approximately 4500 m thick succession within the east greenland caledonian fold belt between latitudes 71°40´ and 74°30´n (haller 1971; peel 1982; henriksen 1985), and were placed in the kong oscar fjord group by smith et al. (2004). the sedimentary rocks laid down in this lower palaeozoic basin are disturbed by large-scale folding and faulting, but are non-metamorphic. lime stones and dolomites dominate the succession which spans the period from the earliest cambrian to the late ordovician (fig. 32); uppermost ordovician to silurian sediments are not known in east greenland (smith & rasmussen 2008). the lower palaeozoic succession begins with c. 200 m of lower cambrian sandstones and siltstones with trace fossils, interpreted as deposited in a tidal to shallow marine environment. these are overlain by a c. 2800 m thick lower cambrian – middle ordovician (dar riwilian) succession of alternating limestones and dolomites, containing a diversified shelf-type pacific fauna (cowie & adams 1957; peel & cowie 1979; peel 1982). stable shelf conditions prevailed throughout the early palaeozoic, with the progressive lithology changes considered to reflect increasing isolation from detrital sources (swett & smit 1972). sedimentary and orga nic–sedimentary structures indicate generally very shallow depositional environments, implying that sed i mentation and subsidence rates were roughly equal. the absence of angular unconformities reflects a non-tectonic environment. the pacific fauna indicates that these areas were developed on the western margin of the proto-atlantic (iapetus) ocean. caledonian orogeny in east and north-east greenland the caledonian fold belts on both sides of the north atlantic developed as a consequence of collision between the continents of laurentia to the west and baltica to the east following closure of the proto-atlantic ocean (iapetus; higgins et al. 2008). the east greenland caledonian fold belt is well exposed between 70° and 81°30´n as a 1300 km long and up to 300 km wide coast-parallel belt. large regions of the fold belt are characterised by reworked precambrian basement rocks [74, 70, 52], overlain by mesoto neoproterozoic [46–43] and lower palaeozoic [40] sedimentary rocks (fig. 33), all of which form parts of westward-directed major thrust sheets. the palaeoproterozoic basement gneiss complexes locally preserve traces of proterozoic fold structures (fig. 34), and have been reworked during caledonian orogenesis. the onshore east greenland caledonian orogen is composed of far-travelled, foreland-propagating thrust sheets that were derived from the laurentian margin and translated westward across the orogenic foreland (higgins & leslie 2000, 2008). the deepest preserved level of the orogen is found north of danmarkshavn (76°40´n) where abundant caledonian eclogitic enclaves occur in the palaeoproterozoic basement gneisses (gilotti 1993; brueckner et al. 1998; gilotti et al. 2008). mediumtemperature high-pressure eclogite facies metamorphism has been dated at 410–390 ma (gilotti et al. 2008). central parts of the orogen (70–75°n) are formed by a pile of several thrust sheets that in their western parts overlie foreland windows (higgins et al. 2004a). the individual thrust sheets include archaean–palaeo proterozoic gneiss complexes overlain by thick successions of latest mesoproterozoic to earliest neoproterozoic sedimentary rocks (krummedal supracrustal sequence). in the highest thrust sheet these sedimentary rocks preserve evidence of early neoproterozoic orogenesis (migmatites and 940–910 ma augen granites). caledonian metamorphism led to emplacement of a suite of caledonian granites, some of which migrated into the basal parts of limestone m 3000 2000 1000 0 middle early late middle early o r d o v ic ia n c a m b r ia n formations heimbjerge narwhale sound antiklinalbugt cape weber dolomite point hyolithus creek bastion kløftelv ella ø dolomite sandstone fig. 32. schematic lithostratigraphic composite section of the cambroordovician sediments in east greenland (c. 71°30´–74°30´n). unit [40] on the map. thicknesses after smith & rasmussen (2008). 51 the overlying eleonore bay supergroup (see also below; kalsbeek et al. 2008b). extensional structures characterise some of the late tectonic phases (strachan 1994; hartz & andresen 1995; andresen et al. 1998; gilotti & mcclelland 2008). the northernmost part of the fold belt in kronprins christian land preserves high-level thin-skinned structures. reviews of the east greenland caledonides have been presented by haller (1971), henriksen & higgins (1976), henriksen (1985) and hig gins et al. (2008). caledonian intrusions and plutonic rocks during the caledonian orogeny widespread migmatisation took place in the crystalline complexes in the southern part of the fold belt and a suite of late to post-kinematic plutons [54] was emplaced in the region between scoresby sund (70°n) and bessel fjord (76°n). north of latitude 72°n the intrusions were emplaced mainly in the boundary zone between the neoproterozoic eleonore bay supergroup sedimentary rocks and the adjacent metamorphic complexes (jepsen & kalsbeek 1998; fig. 35), whereas in southern areas plutonic bodies are widespread within the crystalline complexes. granodiorites and granites are the most abundant types and have yielded intrusive ages ranging from 466 ma to c. 420 ma. most ages occur in the range 440–425 ma (kalsbeek et al. 2008b). the caledonian granites in the northernmost part of their region of occurrence (75–76°n) were emplaced about 430–425 ma ago and these contain a large proportion of crustally derived components (strachan et al. 2001). the southernmost known ‘caledonian’ intrusion is the batbjerg complex (brooks et al. 1981) which occurs in a late archaean granulite facies terrain at kangerlussuaq 68°40´n, c. 200 km south of the nearest exposed part of the caledonian orogenic belt. the batbjerg complex consists largely of pyroxenites including some leucitebearing types [60], and has been dated at c. 440 ma (brooks et al. 1976). major thrust foreland east greenland caledonian fold belt orogen (n ot to s ca le ) ma (windows south of 76°n) north of 78°n south of 78°n franklinian basin (north greenland) hagen fjord group zig-zag dal basalt fm. precambrian shield (incl. metased. + volc.) precambrian shield w in do w s ellesmerian ? (north greenland) caledonian caledonian franklinian basin hagen fjord group rivieradal group independence fjord group basement gneisses t hi n sk in ne d th ru st b el t p ar au to ch . th ru st b el t u pp er th ru st u ni t old red basin caledonian granites basement basement u pp er th ru st s he et kong oscar fjord group tillite group eleonore bay supergroup cambrian ordovician silurian carboniferous devonian 416 542 1000 1600 2500 palaeomesoneop ro te ro zo ic p al ae oz oi c augen granites krummedal sed. grenvil l ian lo w er th ru st s he et n ig gl i s pi ds u ni t h ag ar b je rg u ni t f ra nz j os ep h un it windows (e. greenland) unconformity archaean orogenydetachment fig. 33. schematic diagram showing lithostratigraphy and structural units with pre-caledonian and caledonian elements occurring in the east greenland caledonian fold belt. devonian continental sediments in east greenland following the caledonian orogeny the extensional collapse of the overthickened crustal welt led to the initiation of devonian sedimentary basins in central east greenland (larsen & bengaard 1991; hartz & andresen 1995; larsen et al. 2008). the devonian sediments unconformably overlie ordovician and older rocks, and the deposits were accommodated by se–nw-oriented dip-slip faulting and are preserved in n–s-trending graben-like structures. the basin fill is of middle and late devonian age [39] and consists of more than 8 km of mainly coarse continental siliciclastic sediments with some volcanic intervals. four lithostratigraphic groups have been established, each corresponding to a tectonostratigraphic basin stage (fig. 36). the stages are separated by unconformities related to tectonic events, which took place both during and after sedimentation (haller 1971; olsen & larsen 1993). each basin stage is composed of several depositional complexes and shows approximately similar drainage patterns, with deposition mainly from the west and north. the earliest deposits (vilddal group, up to 2500 m thick) are interpreted as laid down by eastwards draining gravelly braided rivers and sandy and silty allu52 gl._fig. 32 nyt_fig.34 bulletin 185_2 25-05-09 fig. 34. major isoclinal fold in reactivated palaeoproterozoic grey orthogneisses, comprising units of darker banded gneisses and lighter coloured more homogeneous granitoid rocks. the earlier structures have been refolded by n–s-trending open folds with steeply inclined axial surfaces. north side of innermost grandjean fjord (c. 75°10´n), north-east greenland; c. 40 km south-west of ardencaple fjord. the cliff is approximately 1200 m high. fig. 35. caledonian granite with large sedimentary xenoliths of the neoproterozoic eleonore bay supergroup. east of petermann bjerg (c. 73°n), north-east greenland. summit about 2100 m high; upper c. 700 m of cliff face shown. 53 vial fans, which gave way to meandering streams and flood plains. the overlying sandstones (kap kolthoff group, up to 2700 m thick) were deposited by extensive coalescing braidplain systems with southward drainage patterns (olsen & larsen 1993; olsen 1993) dominated by sandy deposits; this group contains intervals of basic and acid volcanic rocks. during the following stage (kap graah group, up to 1300 m thick) sedimentation was dominated by aeolian deposits succeeded by an alluvial and aeolian complex, which was dominated by finegrained sandstones and siltstones. the following stage (celsius bjerg group, 1550 m thick) is characterised by northwards drainage patterns. the deposits comprise fluvial sandstones, flood basin sediments and lacustrine siltstone. the uppermost c. 200 m thick part of the continental celsius bjerg stage is now considered early tournaisian (carboniferous) in age (marshall et al. 1999). celsius bjerg gr. harder bjerg fm. trail ø gr. kap graah gr. kap kolthoff gr. vilddal gr. caledonian folded rocks early midddle late late middle 0 2000 4000 6000 8000 10 000 m d e v o n ia n c a r b o n if e r o u s siltstone conglomerate unconformity sandstone with pebbles fig. 36. schematic, lithostratigraphic composite section of devonian – lower permian continental clastic deposits in central east greenland (71–74°30´n), units [39] and [38] on the map. compiled from olsen & larsen (1993), stemmerik et al. (1993) and p.-h. larsen et al. (2008). north greenland rift basinwandel sea strike-slip mobile belt late cretaceous local pull-apart basins late jurassic _ early cretaceous two separate basins trolle land group late permian _ middle triassic mallemuk mountain group late carboniferous _ early permian sortebakker formation early carboniferousfaultthrust paleogene kap washington group volcanics earliest paleocene peary land kap washington wandel hav trolle land fault zone inland ice kr on pr in s c hr is tia n la nd 75° 50° 25° 82100 km harder fjord fault zone fig. 37. distribution of the wandel sea basin sequences in central and eastern north greenland. modified from håkansson et al. (1994). carboniferous–tertiary deposits of the wandel sea basin, central and eastern north greenland the wandel sea basin deposits were laid down along the northern and north-eastern margin of the greenland shield (figs 37, 38). three main phases of basin formation are recognised, commencing with a widespread carboniferous to triassic event of block faulting and regional subsidence (stemmerik 2000). later, during the late jurassic and cretaceous, more localised basin formation took place during two separate events in a strikeslip zone formed at the plate boundary between greenland and svalbard (håkansson & stemmerik 1989, 1995). lower carboniferous fluvial deposits (sortebak ker formation [20]) are restricted to an isolated halfgraben in southern holm land (c. 80°n; stemmerik & håkansson 1989, 1991). after mid-carboniferous re gional uplift, rifting started in the late carboniferous and more than 1100 m of upper carboniferous to lower permian shallow marine sediments were deposited (mallemuk mountain group [19]) (stemmerik et al. 1996, 1998, 2000). the carboniferous succession is dominated by cyclicly interbedded shelf carbonates (with minor reefs) and siliciclastic rocks. the lower permian is mainly represented by shelf carbonates. renewed subsidence took place during the mid-permian, and the upper permian succession is dominated by alternating shallow marine carbonates and sandstones and deepwater shales. a low-angle unconformity separates these deposits from the overlying lower and middle triassic shelf sandstones and shales (trolle land group [18]) in eastern peary land. sedimentation resumed in the late jurassic, and during the late jurassic and early cretaceous shelf sandstones and shales (ladegårdsåen formation [17]) were deposited in a series of small isolated sub-basins (håkansson et al. 1991). following a new episode of strike-slip movements, renewed sedimentation took place in six minor pull-apart basins during the late cretaceous. each basin is characterised by high sedimentation rates, a restricted lateral extent and its location along strike-slip fault zones (håkansson & pedersen 1982; birkelund & håkansson 1983). depositional environments range from deltaic to fully marine. at kap washington, on the north coast of greenland, c. 5 km of extrusive volcanic rocks and volcanogenic sediments (kap washington group [16]) of peralkaline affinity are preserved (fig. 37; brown et al. 1987). they are of earliest paleocene age (64 ± 3 ma, estrada et al. 2001), and their extrusion may be associated with intrusion of a dense swarm of alkali dolerite dykes in north 54 • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • 4000 2000 0 6000 h: hiatus f: folding sortebakker fm mallemuk mountain group caledonian folded rocks trolle land group thyra ø fm herlufsholm strand fm + correlatives seven informal formations at kilen ladegårdsåen fm c a r b o n if e r o u s p e r m ia n t r ia s c r e t a c e o u s p a le o g e n e ju r m late late late early early early early middle late early major unconformity conglomerate limestone sandstone siltstone dolomite mudstone f f h f h h f fig. 38. composite section of the wandel sea basin successions in eastern north greenland. the successions occur in several distinct sub-basins. corresponding units on the map: sortebakker formation [20]; mallemuk mountain group [19]; trolle land group [18]; ladegårdsåen formation and correlatives [17]; herlufsholm strand formation and correlatives [15]; thyra ø formation [14]. compiled from håkansson & stemmerik (1989), stemmerik & håkansson (1989) and stemmerik et al. (2000). 55 greenland (see fig. 20). the volcanic rocks are preserved below a major, southward-dipping thrust which transported folded lower palaeozoic rocks northwards over the volcanic successions (see fig. 31). all pre-upper cretaceous deposits in eastern north greenland were subjected to compressional deformation during the so-called ‘kronprins christian land orogeny’ (håkansson et al. 1991). subsequently to this deformation event a thin succession of upper paleocene to lower eocene fluviatile and marine sandstones (thyra ø for mation [14]) accumulated, which are the youngest deposits of the wandel sea basin succession (håkansson et al. 1991; lyck & stemmerik 2000). late palaeozoic and mesozoic rift basins in east greenland a series of carboniferous–mesozoic sedimentary basins developed in east greenland following initial postcaledonian devonian deposition. the basins formed as n–s-trending coast-parallel depocentres which reflect prolonged subsidence. important phases of block faulting and rifting took place during the early and late carboniferous, late permian, late jurassic and creta ceous, presaging the opening of the north atlantic in the late paleocene (surlyk 1990, 2003; stemmerik et al. 1993; surlyk & ineson 2003). there is a marked difference in post-carboniferous structural style and depositional history between the basins south and north of kong oscar fjord (c. 72°n). the jameson land basin to the south developed as a late permian – mesozoic sag basin, while the region to the north was characterised by continued block faulting and rifting (fig. 39). initial rifting took place during the latest devonian to earliest carboniferous, when fluvial sandstones and shales were deposited in narrow half-grabens [38] (stemmerik et al. 1991). a pronounced hiatus marked by non-deposition and erosion occurred during the midcarboniferous, and active deposition did not resume until the late carboniferous when up to 3000 m of fluvial and lacustrine sediments were deposited in active half-grabens [38]. deposition ceased sometime during the latest carboniferous or earliest permian. during the early permian a new episode of regional up lift and erosion took place. late permian – early cretaceous deposits of the jameson land basin (70–72°n) the jameson land basin contains a stratigraphically complete succession of upper permian to earliest creta ceous sediments (fig. 40). sediment infill was derived km 0 2 4 0 10 20 30 40 50 w e middle _ late jurassic caledonian crystalline carbonifero s earl permian late permian j rassicd i 0 20 10 20 0 40 60 80 100 120 140 160 km km nw se late jurassic _ early cretaceous 60 km palaeogene basaltcaledonian crystalline basement a b fig. 39. upper palaeozoic – mesozoic basins in east greenland. a: northern development at wollaston forland (c. 74°30´n). b: southern development at jameson land (c. 71°n). note the different scales of the two profiles. from christiansen et al. (1991) and surlyk (1991). 56 from both the east and west during most of the basin history. the first marine incursion into the area since the early palaeozoic took place during the late permian and earliest triassic with deposition of more than 900 m of shallow marine sediments [37] (surlyk et al. 1986). the permian deposits include alluvial fan conglomerates to marginal marine carbonates and evaporites in the lower part, and carbonate platform to basinal shale deposits in the upper part. the latest permian and triassic deposits were dominated by marine sandstones and shales. the next stage in basin development began with deposition of c. 1400 m of alluvial conglomerates and lacustrine dolomite and shale during the triassic [36] (clemmensen 1980a, b). this late palaeozoic – mesozoic extensional basin in east greenland contains a succession from uppermost trias sic to lower cretaceous, recording at first thermal subsidence, then onset and culmination of rifting, and finally waning of rifting (surlyk 2003). a major lacustrine basin [35] covered most of jameson land during the latest triassic – earliest jurassic (dam & surlyk 1993, 1998). renewed marine incursions took place during the early jurassic (dam & surlyk 1998), and during the remaining part of the jurassic and earliest cretaceous shelf conditions persisted in the basin (surlyk 1990). during middle and late jurassic time the basin infill mainly comprised shallow-water sandstones in the northern half of the basin while deeper water shales occur in the southern part [34]. latest jurassic and earliest cre taceous deposits [33] are restricted to the southernmost part of the basin and are dominated by shallow marine sandstones (surlyk 1991). a revised stratigraphy of the uppermost triassic to jurassic has been proposed by surlyk (2003). unconformityconglomerate limestone sandstone siltstone gypsum mudstone hesteelv fm raukelv fm hareelv fm olympen fm vardekløft fm neill klinter gp kap stewart gp fleming fjord fm gipsdalen fm pingo dal fm wordie creek fm foldvik creek gp late late late late middle middle middle early early early p e r m . t r ia s s ic c r e t . ju r a s s ic ju r a s s ic c r e t a c e o u s jameson land wollaston forland m m 4000 3000 2000 2000 1000 1000 0 0 albian_aptian lindemans bugt fm palnatokes bjerg fm bernbjerg fm vardekløft fm caledonian folded rocks • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • fig. 40. schematic sections of the southern (jameson land) and northern (wollaston forland) developments in the late permian and mesozoic rift margin basins of east greenland. corresponding units on the map: foldvik creek group and wordie creek formation [37]; pingo dal, gipsdalen and fleming fjord formations [36]; kap stewart group and neill klinter group [35]; vardekløft, olympen, hareelv and bernbjerg formations and correlatives [34]; raukelv, hesteelv, lindemans bugt and palnatokes bjerg formations and aptian–albian sediments [33]. compiled from: surlyk & clemmensen (1975); clemmensen (1980b); surlyk et al. (1981, 1986); surlyk (1990, 1991); stemmerik et al. (1993); dam & surlyk (1998). a revised provisional lithostratigraphy of the uppermost triassic – jurassic has been proposed by surlyk (2003). late permian – cretaceous deposits in north-east greenland (72–76°n) the sedimentary succession is stratigraphically less complete in this part of east greenland due to continuous block faulting during the mesozoic (surlyk 1990; stemmerik et al. 1993). major hiatuses occur at around the permian–triassic boundary and in the triassic and early jurassic. the upper permian and lower triassic sediments [37, 36] resemble those in jameson land; continental middle triassic deposits are restricted to the southernmost part of the region. the middle to upper jurassic sediments [34] also resemble those in jameson land (fig. 40), but were deposited in a separate basin (surlyk 1977). renewed rifting disrupted the northern part of the region into a series of 10–40 km wide half-grabens during the latest jurassic and earliest cretaceous (surlyk 1978). these were infilled with more than 3000 m of syn-sedimentary marine breccias and conglomerates that pass laterally into sandstones and shales. the younger cretaceous sediments (upper part of [33]) were deposited in a less active rift setting and are dominated by sandy shales with minor conglomerates. cretaceous–palaeogene deposits central west greenland cretaceous–palaeogene sedimentary rocks [8] crop out in the disko–svartenhuk halvø region (69–72°n) of west greenland, where they are overlain by palaeogene basalts (pedersen et al. 2006). the sediments were laid down in the nuussuaq basin and are referred to as the nuussuaq group. although now bounded to the east by an extensional fault system, the sediments may originally have extended both east and south of their present area of outcrop (chalmers et al. 1999). a single seismic reflection line acquired on the south coast of nuussuaq c. 25 km west of locality 2 in fig. 41 suggests that there are at least 6 km, and perhaps as much as 8 km, of mesozoic sedimentary rocks below sea level at this locality (christiansen et al. 1995; chalmers et al. 1999), but the age and character of the deepest deposits are not known. the following notes are drawn largely from a new comprehensive description of the entire succession of exposed and drilled sedimentary rocks in the basin (dam et al. 2009). the oldest sedimentary rocks exposed in the nuussuaq basin belong to the kome formation of albian age (column and locality 4 in fig. 41). these were deposited during a syn-rift phase and crop out on north-east nuus suaq where they lie directly on weathered precambrian basement. the coarse sandstones, mudstones and sparse coal of the kome formation reflect an environment dominated by fluvial channels, flood plains and fan deltas amid basement highs. the slibestensfjeldet formation that overlies the kome formation on north-eastern nuussuaq is up to 240 m thick and was deposited in an extensive lake. in latest albian to early campanian time the southeastern part of the nuussuaq basin was the site of a major fluvio-deltaic system that fanned out to the west and north-west from a point somewhere east of the island of disko/qeqertarsuaq, reaching deeper water at a shelf edge situated approximately at the position of the n–strending fault system crossing disko and nuussuaq that is shown in the inset map in fig. 41. the deposits of this system constitute the atane formation which is at least 3000 m thick in the vaigat area, although the thickest exposed sections are only up to 800 m thick. in the south-east, sandstones alternating with mudstones, coal seams and heteroliths represent the deposits of a braided river plain while farther north-west (e.g. at locality 2 in fig. 41) the formation consists of stacked, typical deltaic, coarsening-upward successions (fig. 42). the coeval deep marine sedimentary rocks on western nuussuaq and eastern svartenhuk halvø are referred to the itilli formation which comprises slope mudstone, turbidite sandstone and conglomerate units (dam & sønderholm 1994). on svartenhuk halvø the formation is dominated by mudstone intercalated with thin beds of sandstone interpreted as distal turbidites (dam 1997). the slope deposits of the itilli formation were also penetrated in the gro#3 exploration well on western nuussuaq (locality 1, fig. 41) that terminated drilling at 3 km depth. the thickness of the itilli formation is estimated to be at least 2000 m on western nuussuaq and at least 1000 m in eastern svartenhuk halvø (dam et al. 2009). in the early campanian the region again became tectonically unstable (dam et al. 2000). phases of blockfaulting and uplift were followed by incision of both submarine and subaerial canyons into the underlying deposits (fig. 42). conglomerates, turbiditic and fluvial sands and mudstones of maastrichtian to danian age (kangilia, quikavsak and agatdal formations) were deposited mainly in valleys and submarine canyons and consequently vary considerably in thickness (dam & sønderholm 1994, 1998; dam 2002; dam et al. 1998, 2009). the kangilia formation varies in thickness from 440 m where it is thickest (locality 3, fig. 41) to only 75 m on central nuussuaq. the quikavsak and agatdal 57 58 thanetian western nuussuaq (gro#3 well) south coast nuussuaq north coast nuussuaq selandian p al eo ce ne u pp er c re ta ce ou s lo w er c re ta ce ou s danian maastrichtian campanian santonian turonian cenomanian albian ages in ma coniacian 99.6 93.5 83.5 70.6 65.5 61.7 v ai ga t s lib es te ns fje ld et a ta ne k om e contact to basement iti lli a ta ne k an gi lia k an gi lia iti lli v ai ga t a ga td al e e q ui ka vs ak t tuff1.5 km undated deep marine sediments m 1 at south coast nuussuaq v ai ga t q ui ka vs ak k an gi lia e m at a ta ne 2 4 3 5 alluvial fan fluvial channel flood plain lake volcanic rocks delta front marine shoreface marine shelf marine slope marine slope channel coal mudstone sandstone conglomerate macrofossils delta plain hyaloclastic breccia subaerial lava flow t t t t estuary palaeogene basalts and intrusives cretaceous-paleocene sediments precambrian basement fault disko 50 km 70° 53° nuussuaq 1 2 5 43 fig. 41. lithostratigraphic sections in the nuussuaq basin, disko–nuussuaq region, central west greenland (from dam et al. 2009). the names of the formations are shown to the left of the simplified logs which indicate the main lithologies and depositional environments. m: maligât formation (belongs to the overlying west greenland basalt group); at: atanikerluk formation; e: eqalulik formation. locations of the sections are shown in the index map. formations are time-equivalents, the former being fluvial–estuarine and the latter marine. the thickness of the entirely channelised quikavsak formation varies from zero to 180 m, while the agatdal formation is 18–65 m thick on central nuussuaq but c. 250 m thick in the gro#3 well. on central nuussuaq some units in the agatdal formation are extremely rich in redeposited marine fossils, mainly gastropods and bivalves. the marine mudstones of the eqalulik formation that overlie the formations mentioned before are locally interspersed with volcaniclastic sandstones and tuffs, thus recording the earliest evidence of volcanic activity within the sedimentary section of the nuussuaq basin. the thickness of the eqalulik formation varies from 12 m to more than 200 m. the youngest deposits (atanikerluk formation) in the nuussuaq basin were deposited in lakes in the eastern part of the basin when basalts in the form of hyaloclastite ‘deltas’ overlain by subaerial lava flows prograded from the west and dammed up a large body of standing water fed by fluvial run-off (g.k. pedersen et al. 1998; a.k. pedersen et al. 2007). the deposits that filled the lakes are arranged in two coarsening-upward succes sions beginning with lacustrine mudstones and passing 59 kf afafafafaafafaafaaaafafaafaa d vfvfvfvffvvfvvqfqfqfqfqfqfqqqfqfqffq vfqf af fig. 42. major submarine canyon incised in middle turonian to late? santonian sediments (atane for mation, af) and infilled with late maastrichtian and early danian turbidites (kangilia formation, kf). thickness of the well-exposed part of the section is approximately 250 m. note the coarsening-upwards cyclicity in the atane formation sediments. pale sandstones of the danian quikavsak formation (qf) can be seen high up on the ridge. the highest rocks exposed are hyaloclastic breccias of the vaigat formation (vf); a dyke is marked d. summit of the ridge is 920 m a.s.l. locality: ataata kuua (locality 2 in fig. 41). up into lacustrine–fluvial sandstones. the cumulative thickness of the deposits is c. 500 m, the thickest single section (c. 400 m) occurring on eastern disko. southern east greenland a c. 1 km thick cretaceous to palaeogene sedimentary succession [50] occurs in east greenland in the kan gerlussuaq basin north-west of nansen fjord (68°17´n). the sediments onlap crystalline basement to the east and north, but elsewhere the base of the succession is not seen. the sedimentary rocks belong to the kanger dlugssuaq and blosseville groups (soper et al. 1976; nielsen et al. 1981). the oldest exposed deposits are fluvial and estuarine sandstones of late aptian – early albian age. they are overlain by upper cretaceous offshore marine mudstones interbedded with thin turbiditic sandstones. in the early paleocene sediment input increased and submarine fan sandstones were deposited along the northern basin margin whereas mudstone deposition continued within the basin. the offshore marine succession is unconformably overlain by fluvial sheet sandstones and conglomerates of latest paleocene age (m. larsen et al. 1999, 2001, 2006). the succession records a basin history of mid-cretaceous transgression and late cretaceous – early paleocene highstand followed by extensive uplift and basin-wide erosion in the mid-paleocene. the uplift was quickly followed by renewed subsidence and the onset of extensive volcanism. tertiary volcanics, intrusions and post-basaltic sedimentary rocks the palaeogene lava regions of both west and east greenland represent major eruption sites at the edges of the continent, from which lavas spilled over mesozoic – early paleocene sedimentary basins and lapped onto the precambrian basement of the continental interior. the volcanic products were formed during the initial phase of continental break-up and initiation of sea-floor spreading in the early palaeogene. palaeogene basalts, central west greenland palaeogene volcanic rocks crop out in central west green land between latitudes c. 69° and 73°n. they are noted for the presence of native iron-bearing basalts and large volumes of high-temperature picrites and olivine basalts (clarke & pedersen 1976; l.m. larsen & pedersen 2009). the composite stratigraphic thickness of the succession varies between 4 and 10 km, with the smallest thickness on disko and a maximum on ubekendt ejland/illorsuit (71°n). eruption of the basalts began in a submarine environment, and the earliest basalts, which occur to the west (fig. 41), consist of hyaloclastite breccias. when the growing volcanic pile became emergent, thin subaerial pahoehoe lava flows started to form. they flowed eastwards into a deep marine embayment where they became transformed into hyaloclastite breccias that prograded eastwards in large-scale gilbert-type deltas with foresets up to 700 m high (pedersen et al. 1993). blocking of the outlet caused the marine embayment to be transformed into a lake, which was ultimately filled in with volcanic rocks (a.k. pedersen et al. 1996; g.k. pedersen et al. 1998) so that subsequent lava flows lapped onto pre cambrian crystalline basement highs in the east. the lower part of the succession (vaigat formation) consists almost entirely of tholeiitic picrites and olivinephyric to aphyric magnesian basalts [7] (pedersen 1985a). the upper part of the succession (maligât formation) consists of tholeiitic, plagioclase-phyric basalts [6] which formed thick plateau lava flows of aa-type. both the vaigat and maligât formations contain sediment-con taminated units of magnesian andesite and, in the maligât formation, also dacite and rhyolite, mostly as tuffs (e.g. pedersen 1985b). some of the sediment-contaminated rocks in both formations contain graphite and native iron, formed by reaction with coal and organic-rich mudstones. the succession is mostly flat lying, but is cut by coast-parallel faults in the western part where the lavas dip at up to 40° westwards. on svartenhuk halvø the upper part of the succession is named the svartenhuk formation, which is the stratigraphic equivalent to the maligât formation (j.g. larsen & pulvertaft 2000). this area is characterised by extensional faulting and tilting, which together with flexure zones locally give rise to dips of up to 60° to the south-west. the major part of the volcanic pile was erupted in a short time span 62–60 ma ago. on western nuussuaq and svartenhuk halvø there is a younger group of lavas dated at c. 55 ma (recognised after the map was printed) which also occurs on ubekendt ejland (storey et al. 1998). 60 61 palaeogene basalts, east greenland early palaeogene volcanic rocks crop out in east greenland between latitudes 68° and c. 75°n. south of scoresby sund/kangertittivaq (c. 70°n) plateau basalts cover an extensive region of c. 65 000 km2, resting on meso zoic–paleocene sediments in the east and south, and on caledonian and precambrian gneisses in the west (nielsen et al. 1981; l.m. larsen et al. 1989). north of scoresby sund palaeogene basic sills and dykes are widespread in the mesozoic strata, and a further sequence of plateau basalts is found between latitudes 73° and 75°n. blosseville kyst region (68–70°n) the earliest palaeogene volcanics are 61–58 ma old (storey et al. 2007). they comprise a c. 1.8–2.5 km thick succession of tholeiitic basalts with subordinate picrite [49], which occurs in the southernmost part of the volcanic province between 68° and 68°30´n (nielsen et al. 1981). the basalts are aphyric or olivine-pyroxene-phyric, and the succession consists of intercalated subaerial flows, hyaloclastites, tuffs and sediments. it is interpreted as the infill of a shallow, partly marine, basin with a source area to the south, along the present coast or on the shelf. the main part of the region 68–70°n is made up of a thick succession of 56.1–55.0 ma old tholeiitic plateau basalts [48] formed by 5–50 m thick subaerial flows of plagioclase-phyric to aphyric basalt (l.m. larsen et al. 1989, 1999; pedersen et al. 1997; storey et al. 2007). the succession is at least 5.5 km thick in the central blosseville kyst area and thins inland and to the north to 2–3 km (fig. 43). four formations can be followed over almost the whole area and represent two major volcanic episodes. eruptions took place over the entire area, but accumulation was largest in the coastal areas where the lava pile sagged during deposition. the subsidence accelerated with time, suggesting increased focusing of the magmas into a developing rift zone beyond the present coast. along the present coast the lava flows dip seawards at 10–50° due to later flexing and faulting (nielsen & brooks 1981; pedersen et al. 1997). intense injection of coast-parallel dykes occurred in several episodes (nielsen 1978; l.m. larsen et al. 1989). younger alkali basalt lavas cap the plateau basalts in some small inland areas; one of these occurrences is of miocene age (13–14 ma; storey et al. 1996). hold with hope to shannon region (73–75°n) a succession of c. 600–800 m of plateau basalts [32] occurs in the hold with hope to shannon region in a block-faulted area. the succession is divided into a lower part of uniform tholeiitic lavas and an upper part with variable tholeiitic and alkali basaltic lavas (upton et al. 1984, 1995; watt 1994). between the two there are local occurrences of intervolcanic conglomerates. the basalts on shannon and the pendulum øer, north-east of wol fig. 43. major unconformity between precambrian gneisses deformed during the caledonian orogeny [52] and palaeogene plateau basalts [32]. the basalt section shown is approximately 800 m thick. north of gåsefjord/nertiit kangersivat (c. 70°n), scoresby sund region, central east greenland. photo: w.s. watt. 62 laston forland, mainly occur as voluminous sills. the reduced magnitude of volcanic activity in this northerly area, compared to the region south of scoresby sund, suggests that it was peripheral to the main volcanic activity in the east greenland tertiary volcanic province. small areas of basalts with alkaline chemistry occur in the nunatak region (74°n) where they overlie caledonian and older crystalline rocks (katz 1952; brooks et al. 1979; bernstein et al. 2000). palaeogene intrusions, east greenland numerous intrusions are exposed along about 1000 km of the coastal region of east greenland between latitudes 66°30´ and 74°n, in addition to the many dykes and sills (see fig. 20); approximately 20 of these intrusions are shown on the map, separated into felsic [53] and intermediate and mafic [57] types (fig. 44). they reflect episodes of alkaline magmatism linked to the continental break up of the north atlantic (nielsen 1987), and range in age from late paleocene to oligocene. the oldest intrusions occur in the south, and have ages between 57 and 47 ma. felsic intrusions in the south are 35–37 ma old, whereas the more northerly intrusions (72–74°n) have ages in the range 48–25 ma (tegner et al. 1998, 2008; brooks et al. 2004). petrologically the intrusions can be divided into three groups (nielsen 1987): (a) alkaline inland intrusions; (b) alkaline dyke swarms and (c) syenitic to granitic complexes and dykes. most of the numerous intrusions found along the coast belong to the third group; they are central intrusions and intrusive complexes, often with several rock types within the same complex. they range from a few square kilometres to c. 850 km2 in size. the felsic complexes [53] are dominated by alkali granites, quartz syenites, syenites and nepheline syenites. the mafic to intermediate complexes [57] are dominated by tholeiitic gabbros, whereas subordinate rock types locally include monzonite and alkali gabbro. the 55 ma old skaergaard intrusion is a classic example of a layered gabbroic intrusion, and has been studied in great detail (wager & deer 1939; mcbirney 1996a, b; irvine et al. 1998). post-basaltic palaeogene sedimentary rocks, east greenland post-basaltic sedimentary rocks [47] are preserved in two small, down-faulted areas near the atlantic coast south of scoresby sund (kap brewster, c. 70°10´n and kap dal ton, c. 69°25´n). they comprise a c. 100 m thick suc ces sion of palaeogene (middle eocene, 48–45 ma old) fluvial to shallow marine sandstones and siltstones (m. larsen et al. 2005) referred to as the kap dalton group. these post-volcanic deposits have been dated by dinoflagelate cysts and suggest that volcanism in this region came to an end close to the early-middle eocene boundary between 49 and 48 ma. the palaeogene sedimentary rocks preserved onshore are marginal exposures of an extensive and much thicker (5–6 km) tertiary succession found on the adjacent shelf areas (see p. 75). pliocene–pleistocene sediments, central north greenland the late pliocene – early pleistocene kap københavn for mation [13] is a c. 100 m thick succession of unconsolidated sand and silt, which crops out over an area of c. 500 km2 in easternmost peary land, north greenland s coresby sund inland ice 27°33° 72° 68° iceland d a n m a r k s t r æ d e felsic intrusions intermediate and mafic intrusions kangerdlugssuaq gustav holm kap kialineq imilik nualik skaergaard kap edvard holm borgtinderne lilloise hold with hope kap parry kap simpson werner bjerge 100 km fig. 44. major tertiary intrusive centres in east greenland (c. 66°30´–74°n). slightly modified from nielsen (1987). (funder & hjort 1980; funder 1989). the succession contains well-preserved faunal and floral elements. the base of the succession is not exposed. the lower 25 m comprise marine silt containing high arctic molluscs, whereas the upper sand-dominated part containing tree trunks reflects nearshore environments. the flora and fauna found in this upper unit point to a much warmer climate than the present. the kap københavn formation shows disturbance caused by overriding glaciers during the quaternary glaciation, and is overlain by till. quaternary glacial sediments and glaciation during most of the quaternary greenland was completely, or almost completely, covered by ice, and glacial deposits are widespread on the present ice-free land areas and on the adjacent shelf (funder 1989; funder et al. 1998). as the map is a bedrock geology map, quaternary deposits are only shown in regions where a thick cover of quaternary superficial deposits conceals the bedrock over large areas (valleys, interior plains and some coastal areas). these areas have been shown on the map as undifferentiated quaternary. recent studies indicate that the glaciation of northeast greenland had started as early as the mid-miocene (14–15 ma ago; thiede et al. 2001). evidence from the shelf areas shows that an early glaciation of greenland at the end of the pliocene (c. 2.4 million years ago) was more extensive than any succeeding glaciation, with an ice sheet covering nearly the entire shelf region up to a few hundred kilometres beyond the present coastline (funder 1989). during this glaciation the land area was subjected to extensive erosion, with much of the eroded material being deposited on the offshore shelves. the superficial deposits found on the ice-free land areas are dominated by the late quaternary development of the past c. 130 000 years (saalian/illinoian–holocene). the last interglacial period (eemian/sangamonian) is recorded in both east and west greenland. during the late weichselian/wisconsinan c. 18 000 years ago the maximum extent of the ice around the northern parts of greenland was close to the present coastline, whereas in parts of west and south-east greenland the ice advanced onto the shelf area (funder & hansen 1996; fig. 45). in south greenland, modelling of the thickness of the ice cover over the outer coast during the last glacial maximum shows that the ice must have been at least 1500 m thick (bennike et al. 2002). recent studies indicate that the greenland ice sheet during the late weichselian/wisconsinan reached out to the middle– outer continental shelf in north-east greenland, a distance of more than 100–200 km beyond the present coastline (evans et al. 2009). the retreat of the inland ice after the last glacial period began 14 000–10 000 years ago, and continued with oscillations to a maximum stage of withdrawal approximately 6000 years ago when the ice margin was up to 20 km inside its present position. the position of the margin of the inland ice where it now abuts against land areas only shows minor fluctuations (fig. 46). significant changes are almost restricted to major drainage outlets where the inland ice flows into fjords to form calving gla ciers; the most active glaciers in greenland have velocities of up to 22 m per 24 hours. 63 30° w 80° 70° 60° 400 km 50° w 40° w ? ? ? ? greenland inland ice fig. 45. former extent of the greenland inland ice during the last glacial maximum. green: c. 18 000 years before present; red: c. 10 000 years ago; blue arrows: major glacier outlet streams. modified from funder & hansen (1996). the present ice cover of greenland is a relic of the pleis tocene ice ages. it consists of the large continental ice sheet (the inland ice), and local ice caps and glaciers (weidick 1995). the inland ice has an area of c. 1 707 000 km2 and reaches an altitude of 3230 m with a maximum thickness of 3420 m. the local ice caps and glaciers cover areas of c. 49 000 km2 (weng 1995). the volume of the inland ice has been estimated at 2 600 000 km3, based on ice thickness measurements by airborne radio-echo sounding; a rough estimate of the volume of local ice caps and glaciers is 20 000 km3. on the map, surface contours, isopachs of ice thickness and contours of the bedrock below the inland ice are shown. mean annual air temperatures on the inland ice range from –30°c over a large region in its central and northern parts to about –5°c in its south-western marginal areas. the temperature of the ice ranges between –32° and 0°c; with increasing depth, the temperature generally increases due to geothermal heat flux and internal heating caused by deformation. in some locations, the temperature at the base of the ice sheet may reach its melting point. mass balance the mass balance (budget) of the inland ice is the difference between accumulation (of snow in the interior region mainly) and ablation by melting and calving of icebergs in the marginal areas. the accumulation of snow decreases from south to north from more than 2000 mm water equivalent/year in coastal areas in the south-west to 100 mm water equiv64 glaciology fig. 46. characteristic front of the inland ice abutting the ice-free land area, with moraines and small lakes. the distance from the bottom of the picture to the land area in the background is approximately 5 km. the locality is about 75 km north-north-east of søndre strømfjord airport, southern west greenland, at c. 67°30´n. view is towards south. photo: h. højmark thomsen. 65 alent/year or less in interior north-eastern areas (ohmura & reeh 1991). melt rates also decrease from south to north. away from the coast in south-west greenland, the annual melting of the ice at sea level probably reaches values near 10 000 mm water equivalent. however, even along the northernmost margins of the inland ice significant melting occurs; melt-rate models predict values near 2000 mm water equivalent/year at sea level. calving glacier fronts producing icebergs are generally located at the heads of fjords at some distance from the outer coast. the most concentrated source region for icebergs is central west greenland (disko bugt and the area between nuussuaq and svartenhuk halvø) where about 100 km3 of calf ice are produced annually. the effects of climate change in recent years on the mass balance of the greenland ice sheet have been documented by satellite gravity measurements. over the four years 2004–2007 the ice sheet lost an average of c. 400–500 km3 in the summer period of each year and only gained c. 250–350 km3 of snow in the winter. the net result is a loss of c. 150 km3/year from the beginning of the 20th century (witze 2008), although some researchers estimate even larger figures for the present net loss. past climate and environment up to 2009 five deep ice cores have been retrieved by drilling through the inland ice (one drilling was only to a depth of 1400 m), and these have provided considerable information about climate and environmental variations during the past 150 000 years. the ice-core records indicate that in central greenland the inland ice survived the last interglacial (the eemian) which culminated about 125 000 years ago, without completely disappearing even when the climate was several degrees warmer than at present. however, according to icedynamic model calculations of the evolution of the inland ice, the ice cover in northern and southern greenland was less extensive during the eemian (fig. 47). the ice-core records indicate dramatic temperature fluctuations during the last ice age, which lasted from about 115 000 years ago to about 11 700 years ago. in the cold est parts of this period, temperatures in greenland may have been 10–12°c colder than now, whereas temperatures in other periods of the ice age were only about 5 degrees colder (dansgaard 1997; hammer 1997). a b c 400 km 130 000 years 15 000 years 0–1000 m 1000–1600 m 1600_2400 m 2400–3000 m > 3000 mice free present fig. 47. models of the inland ice with indication of thickness of the ice sheet in metres. a: the last interglacial (the eemian) with a tem perature 4–5°c higher than the present. b: during the late glacial maximum (weichselian) with a temperature 10–12°c colder than at present. c: under the present climatic conditions. from model calculations by letréguilly et al. (1991). the models do not include the offshore extent of the ice, only that of present land areas. the interpretation of the offshore geology shown on the 1:2 500 000 scale map was based mainly on seismic surveys carried out between 1970 and 1992, supplemented by aeromagnetic and gravimetric data and, in the case of southern west greenland, by data from five exploration wells drilled in 1976–77. offshore south-east greenland six holes were drilled in 1993 at c. 63°n, constituting leg 152 of the ocean drilling program (odp; h.c. larsen et al. 1994a); the positions of three of these wells are shown on the map. in 1995 three more holes were drilled off south-east greenland as part of the aborted odp leg 163, but no results were available when the map went to press (h.c. larsen et al. 1996). however, the coverage of geophysical data in different areas was, and still is, uneven, and is dependent on ice conditions. off southern west greenland, where there are only scattered icebergs and no pack ice in the late summer and early autumn, more than 37 000 km of seismic data were acquired by the oil industry in the 1970s and a further 10 259 km of non-exclusive data were acquired in this area in the years 1990–1994 (pulvertaft 1997). in contrast, the often ice-infested areas off east, north-east and north-west greenland were only covered by reconnaissance surveys, principally as a result of the kanumas and north atlantic d (nad) surveys. the kanu mas project was a marine seismic reconnaissance financed by six major oil companies, with the greenland-danish national oil company nunaoil a/s as operator (h.c. larsen & pulvertaft 1990; pulvertaft 1997). in the 1990s kanumas surveys acquired c. 7000 km of seismic data off north-east and central east greenland, and c. 4000 km of data off north-west greenland; although these data are confidential company data, some results had been released in time to be included in the 1:2 500 000 map. the north atlantic d project was a combined aeromagnetic and seismic survey of the east greenland shelf carried out by grønlands geologiske undersøgelse (ggu) in 1979–83. during this project c. 8000 km of seismic data were acquired off central east and southeast greenland (thorning et al. 1982; h.c. larsen 1985). in nares stræde (nares strait) and off north greenland, where no seismic data existed, interpretation of the geology was based on aeromagnetic and sparse gravity data alone. aeromagnetic and shipborne magnetic data constituted the main source of information in oceanic areas. the 1:2 500 000 map was designed to show two general aspects of offshore geology: (1) the extent of continental crust [a], oceanic crust [c–g], and of the intervening, poorly understood, transition zone [b] and (2) the distribution of sedimentary basins and major faults. where extensive volcanic units are known to occur in areas underlain by continental crust, their distribution is also shown. since compilation of the 1:2 500 000 map a large amount of new data has been acquired in the maritime regions surrounding greenland, both by the industry and by academic research institutes. it is clear from these data that the map is not correct in many places. in the following text, attention is drawn to the known errors in the map, and as more data are released, there will no doubt be need for further corrections. the distribution of crustal types offshore as now understood (2009) is shown in fig. 49a, p. 68, while offshore and onshore sedimentary basins are shown in fig. 56. the continental margin off east and north greenland east greenland south of 77°n in general terms, the continental margin off east greenland between the southern tip of greenland and 76°n can be described as a volcanic rifted margin (h.c. larsen 1990; h.c. larsen et al. 1994a, b), formed when greenland became separated from northern europe at the start of sea-floor spreading in early eocene time (magnetochron 24r). between c. 68°n and the jan mayen fracture zone, however, greenland remained attached to the jan mayen microcontinent (fig. 49a, b) until oligocene time when spreading shifted from the aegir ridge to the kolbeinsey ridge. the position of the continent–ocean boundary (cob) was drawn on the basis of aeromagnetic data supplemented by characteristic features in the nad reflection seismic data. the absolute seawards (eastern) limit of continental crust cannot overlap areas where linear magnetic anomalies characteristic of oceanic crust can be identified with confidence. along the entire volcanic rifted margin seaward-dipping reflectors can be seen in the seismic data. these arise from subaerial lava flows or groups of flows that were erupted in the early stages of 66 offshore geology sea-floor spreading prior to differential subsidence below sea-level. in connection with the seaward-dipping reflectors, buried volcanic escarpments may occur. these are landward-facing escarpments formed at the landward end of the dipping reflectors, where lava flows interfinger with sedimentary rocks (fig. 48; h.c. larsen & jakobsdóttir 1988; h.c. larsen 1990). the zone off east greenland shown on the map as underlain by transitional crust [b] was drawn in a rather arbitrary manner, at least with regards to its width. this zone is thought to consist of attenuated and fragmented continental crust with increasing numbers of dykes and other intrusions as oceanic crust is approached. much of the onshore coastal area around and south of kang erlussuaq (68°n) is very intensely intruded by palaeogene coast-parallel dyke swarms (nielsen 1978; klausen & larsen 2002; not shown on the map). aeromagnetic data indicate that these dyke swarms continue southwestwards under the shelf as far south as 63°n (h.c. larsen 1978). this intensity of dyke intrusion suggests the proximity of the continent–ocean boundary, i.e. the outer edge of the transition zone. since the map was printed, results of intensive research carried out off southern east greenland along leg 152 of the ocean drilling program (odp sites 914–919, c. 63°n) have been published (h.c. larsen & saunders 1998; h.c. larsen et al. 1998). results of this research indicate that the continent–ocean boundary, defined here as the point at which thinned, intensely dyked continental crust finally gives way to a sheeted dyke complex, is situated in this area about 12 km landwards of the shelf break (h.c. larsen & saunders 1998 fig. 12). the shelf break here is the edge of a thick prograding wedge of glaciomarine sediments. the inner boundary of the continent–ocean transition zone, i.e. the point at which extensional faulting intensifies and marked attenuation of continental crust begins, lies 25–40 km landwards of the continent–ocean boundary (h.c. larsen & saunders 1998 fig. 12; h.c. larsen et al. 1998 fig. 7). thus the continent–ocean transition zone may be a little wider than shown on the 1:2 500 000 map, and the continent–ocean boundary probably lies about 25 km north-west (landwards) of the position shown on the map. at c. 68°n the eastern margin of continental greenland cuts obliquely across linear magnetic anomalies 24–13 in the oceanic crust. this was not originally regarded as the expression of a transform fault, but rather as an oblique ocean–continent transition along a former northward-propagating spreading ridge (h.c. larsen 1988). however, prior to anomaly 13 time, the jan mayen mi cro continent was attached to east greenland between c. 68° and 72°n. a coast-parallel dyke swarm along blosseville kyst between 68°20´ and 70°n and voluminous intraplate volcanism in this region may reflect an unsuccessful attempt at continued eocene spreading along an axis at about the position of the present coast (h.c. larsen 1988). however, to find the ‘missing’ magnetic anomalies 24–13 (i.e. eocene) oceanic crust, it is to the east of the jan mayen micrcontinent that one should look (e.g. lundin & doré 2002). during this period a transform fault must have linked the reykjanes ridge to the southern end of the aegir ridge – the denmark strait fracture zone (lundin & doré 2002). after anomaly 13 time spreading between greenland and the jan mayen microcontinent propagated northwards, reaching the jan mayen fracture zone at about anomaly 6 time. north of jan mayen fracture zone the 1:2 500 000 map shows the cob off east greenland transgressing magnetic anomalies 24b, 24a and 23 at a low angle, indicating that here the spreading ridge propagated towards the south-west. newer interpretations of the position of the cob here differ, not only from what is shown on the map but also from one another. tsikalas et al. (2002) extend anomalies 24b and 24a into the shelf, the anomalies crossing the shelf edge at approximately 74°15´n and 73°55´n respectively. this implies that the shelf has prograded over oceanic crust here, and that there was no south-westwards progradation of the spreading axis in this region. however, refraction seismic data from the shelf between 72° and 74°n (voss & jokat 2007) show clearly that the boundary of true oceanic crust lies very slightly seawards of what is shown in the 1:2 500 000 map, while the transition zone, described by voss & jokat as “intruded and stretched continental crust”, extends landwards 100 km from the cob. farther 67 isochrones basalts direction of lava flow 2 km sediments sediment flow pse pse pse a pp ro x. 1 km fig. 48. cross-section based on seismic section, illustrating the formation of so-called pseudo-escarpments (pse) at the landward end of dipping basalts. sediments: pale blue to light brown layers; basalts: green. landward direction to the left. slightly modified from h.c. larsen (1990). 68 g r e e n l a n d s v a l b a r d e l l e s m e r e i s l a n d b a f f i n i s la n d i c e l a n d 75 ° 39 ° 3°w 9°e 82 ° 78 °86° 82° 74° 70° 66° 62° 63 ° 58° 33° 21° 66° 3° 62° 15° 74 ° 39°51°54° 58° 99° 111° 0 250 500 km qh gr a north berger & jokat (2008), on the basis of reflection seismic data, place the landward boundary of oceanic crust a little seawards of the position shown on the enclosed 1:2 500 000 map, the difference between berger & jokat’s interpretation and the 1:2 500 000 map diminishing south-westwards. north-east greenland (77–83°n) at the greenland ridge (see fig. 49a, b) a sliver of continental crust extends from the shelf south-eastwards along the fracture zone until this turns slightly anticlockwise. north-east of the fracture zone there is an area tentatively interpreted as extremely thinned continental crust (døssing et al. 2008); this is included in the area shown as continental or transitional crust in fig. 49a. north of the fracture zone the position of the cob steps to the north-west. sea-floor spreading magnetic anomalies have lower amplitudes and are more ambiguous in the oceanic area north of the fracture zone and in the fram strait than in the oceanic region to the south. at the time the 1:2 500 000 map was compiled, there was no agreed interpretation of magnetic anomalies, and hence 69 ridge g.r.g.r. e l l e s m e r e e l l e s m e r e i s l a n d i s l a n d b a f f i n i s la n d n.w.bn.w.b. „ 75 ° 82 ° 78 ° 99° 86° 74° 70° 66° 62° 63 ° 58° 111° 33° 21° 66° 62° 58° 15° 74 ° 70 ° 39° 51° 54° g r e e n l a n d l i n c o l n s e a e u r a s i a b a s i n e l l e s m e r e i s l a n d m.j .r . g.r. s.f.z. m.f.z. y.p l . lo.r. f r a m s t r a i t g ak ke l s v a l b a r d r ey kj an es r id ge k o lb e in se y r id ge j an mayen f.z. m o hn s r id ge r idge kn ipov ich l a b r a d o r s e a ? u .f. z. d. s.h . b a f f i n i s la n d b.p. c.d. jan m ayen m c. s .sd . n ares s t . k .k . r id ge 82° 39 ° 3°w 15° 9°e n.w.b. „„ b a f f i n b a y i c e l a n d 250 500 km0 b land areas. maritime areas underlain by continental or transitional crust oceanic crust, pliocene–recent oceanic crust, miocene oceanic crust, oligocene oceanic crust, eocene oceanic crust, paleocene oceanic crust, age uncertain palaeogene basalts exposed at seabed in areas underlain by continental or transitional crust palaeogene basalts buried under younger sediments in areas underlain by continental or transitional crust seaward-dipping reflectors oceanic spreading axis, active oceanic spreading axis, extinct transform fault major basin – bounding fault; mark on downthrown side compressional fault, thrust exploration well; q: qulleq-1, g: gjoa g-37, r: ralegh n-18, h: hekja 0-71 fig. 49. a (facing page): map showing distribution of crustal types and major structures offshore greenland. only basalts overlying continental crust are shown. b (above): place names and names of structural features shown in a or mentioned in the text. lo.r.: lomonosov ridge; m.j.r.: morris jesup rise; y.pl.: yermak plateau; s.f.z.: spitsbergen fracture zone; m.f.z.: molloy fracture zone; g.r.: greenland ridge; jan mayen f.z.: jan mayen fracture zone; jan mayen mc.: jan mayen microcontinent; u.f.z.: ungava fracture zone; d.s.h.: davis strait high; ‘n.w.b.’: ‘north water bay’; s.sd.: smith sound; k.k.: kennedy kanal; b.p. bache peninsula; c.d.: cape dyer. legend, facing page: the age of oceanic crust between the hovgård ridge (position c. 0°w, 78°30´n, shown as transitional crust) and the spitsbergen fracture zone (see fig. 49a, b) is shown as [g] (age unspecified) on the map. recently, however, engen et al. (2008) have reviewed the geology and evolution of the fram strait and provisionally identified and numbered several sea-floor spreading magnetic anomalies in the seaway, so that the oceanic crust in the strait can now be subdivided chronostratigraphically as it is elsewhere (fig. 49a). between the greenland ridge and the molloy fracture zone (fig. 49a, b) the oldest sea-floor spreading magnetic anomaly that can be identified with confidence is anomaly 18 (39 ma; middle eocene) (engen et al. 2008). regarding the cob here, døssing et al. (2008) place the seaward limit of continental crust slightly east of where it is shown on the 1:2 500 000 map, close to the shelf edge. the north-east margin of continental greenland, north of 79°n, has a very different character. it is shown on the enclosed map as a former intracontinental transform plate boundary. a consensus exists that a substantial dextral displacement of svalbard relative to north-east greenland took place along an intracontinental nw–se megashear, the de geer megashear (harland 1969; eng en et al. 2008), in the time interval corresponding to mag netochrons 24r–13 (earliest eocene – earliest oligocene). in the early oligocene, rifting began to take place along this zone, and a spreading ridge, the knipovich ridge, linking the mohns ridge and the gakkel ridge, developed along the site of the earlier transform fracture. since this time, the flanking continental margins have developed as passive margins separated by an obliquely spreading ocean (e.g. vogt & tucholke 1989; eldholm et al. 1990; kristoffersen 1990; engen et al. 2008; faleide et al. 2008). spreading propagated south-eastwards from the gakkel ridge so that the earliest identified magnetic anomaly between the yermak plateau and the morris jesup rise is anomaly 7, while in the fram strait the old est anomaly approaching the spitsbergen fracture zone is anomaly 5 (engen et al. 2008). it has also been suggested that mantle peridotite has been exposed in this part of the strait (jokat et al. 2005). whatever the case, the 1:2 500 000 map is incorrect with regard to the age of the oceanic crust in this area. the morris jesup rise and the yermak plateau one particular outstanding problem here is the nature of the crust underlying the morris jesup rise and its conjugate feature, the yermak plateau, north of svalbard (fig. 49a, b). recent work has shown that the yermak plateau south of 82°n (the area shown without colour on the 1:2 500 000 map) is most likely underlain by continental crust, and that the south-west margin of this part of the plateau is a continental margin (ritzmann & jokat 2003; engen et al. 2008). north of 82°n the yermak plateau runs ne–sw; it is this outer part of the plateau that adjoined the morris jesup rise prior to the opening of the eurasia basin and fram strait. the nature of the crust in the outer yermak plateau is still open to discussion. jokat et al. (2008) favour a model of stretched and intruded continental crust, and engen et al. (2008) argue against the earlier hypothesis that this part of the plateau formed by voluminous oceanic volcanism at an eocene triple junction between eurasia, greenland and north america, pointing out that the termination of sea-floor spreading magnetic anomalies 23–13 at the north-east flank of the outer plateau favours a continental outer yermak plateau. a limited amount of work has been done on the conjugate feature, the morris jesup rise. available data indicate that the feature is underlain by thinned and rifted continental crust (ostenso & wold 1977). as is also the case in the yermak plateau, high-amplitude and irregular magnetic anomalies suggest the presence of volcanic rocks, possibly of cretaceous age (engen et al. 2008). dawes (1990) favoured the view that the morris jesup rise has a complex structure and that it may contain appreciable continental remnants below a thick cover of volcanic rocks. the continental margin off west greenland southern west greenland (58–64°n) the continental margin off southern west greenland also presents problems of interpretation, although more reflection seismic data are available from this area. there has also been some controversy as to when sea-floor spreading in the labrador sea began, as discussed in the following, but all agree that sea-floor spreading between greenland and canada began earlier than in the north atlantic, and that it had ceased by magnetochron 13 time. distinct linear magnetic anomalies can be seen in the labrador sea off south-west greenland (srivastava 1978). the earliest magnetic anomalies trend nnw–sse, while the younger anomalies (24 and younger) trend 70 nw–se, parallel to an extinct spreading axis roughly midway between canada and greenland. the oldest unambiguous magnetic anomaly in the labrador sea is anomaly 27. the crust landward of this is shown as transitional crust on the 1:2 500 000 map. srivastava (1978) and roest & srivastava (1989) indicated that linear magnetic anomalies could be identified much closer to the continental shelf than anomaly 27, suggesting 31 or 33 as the number of the oldest anomaly. however, a substantial body of evidence has accumulated since 1991 that refutes the hypothesis that pre-anomaly 27r oceanic crust exists in the labrador sea. modelling of the magnetic data acquired by the bundesanstalt für geowissenschaften und rohstoffe (bgr), during seismic transects across the labrador sea in 1977, indicates oceanic crust with alternating strips of normally and reversed magnetised crust landwards as far as anomaly 26r or 27r (chalmers 1991; chalmers & laursen 1995). landwards of this, a model assuming thinned and rifted continental crust intruded by reversed magnetised igneous material provides the best fit with the observed data. south of 62°n it appears that serpentinised peridotite subcrops sediments in the outer part of the transitional zone and that the remnants of continental crust are very thin (chian & louden 1994; chalmers 1997). whatever the case, the seawards limit of normal continental crust off southern west greenland lies well to the south-west of the continental slope, at water depths of more than 1500 m. this interpretation of the distribution of crustal types is supported by the structural pattern seen in the seismic lines. both the normal continental crust and the zone of transitional crust show large tilted fault blocks overlain by synand post-rift sediments (chalmers & pulvertaft 2001). the oldest sediments are most likely of early cretaceous age (see later). definitive evidence concerning the crustal structure in the inner labrador sea was obtained when the qulleq-1 exploration well was drilled in 2000 at 63°49´n, 54°27´w. according to srivastava & roest (1999 fig. 5), this should lie on magnetic anomaly 33 (73.6–79 ma, middle cam panian; cande & kent 1995), but the well terminated in santonian sandstone without reaching basement (christiansen et al. 2001), showing that the basement here must be older than oceanic crust anywhere in the north atlantic region north of 56°n. the crustal structure in the inner labrador sea has recently been studied in a reflection/refraction seismic transect running wsw from the qulleq-1 well to the gjoa g-37 and hekja 0-71 wells off eastern canada. this new seismic line shows that oceanic crust is restricted to two narrow zones at or close to the ungava fracture zone, a major transform zone that transferred sea-floor spreading through the davis stræde (davis strait) into baffin bugt (baffin bay; fig. 49a, b; funck et al. 2007). funck et al. believe that these zones of oceanic crust formed when phases of transtension along the fracture zone created gaps which were filled with melt that formed new oceanic crust. davis stræde (c. 64–69°n) the crust under davis stræde is estimated to be 22 km thick (keen & barrett 1972), which is intermediate between the thickness of normal oceanic and continental crust. there is a basement high within the strait (the davis strait high), where the sedimentary cover is thin and locally virtually absent. palaeogene volcanic rocks have been recovered from the high (srivastava 1983; williamson et al. 2003), and it has been suggested that the high is a volcanic plateau formed by hotspot volcanic activity. hotspot activity is also suggested by the occurrence of thick picritic lavas of paleocene age in the disko – nuussuaq – svartenhuk halvø area in central west greenland and at cape dyer (fig. 49b) on the south-east side of baffin island in canada (clarke & upton 1971; clarke & pedersen 1976). on the greenland side of the strait, strata of late cretaceous age have been traced on seismic lines westwards from the ikermiut-1 exploration well (66°56´n) onto the eastern flank of the high (chalmers et al. 1995), indicating that the high is formed of pre-late cretaceous rocks. an e–w seismic profile across the high west of the ikermiut-1 well shows dipping internal reflectors within the high (gregersen & bidstrup 2008), suggesting the presence a sedimentary unit. the very high abundance of ordovician carbonates in dredge samples collected on the high suggests that this unit consists of ordovician limestones (dalhoff et al. 2006), a possibility supported by the occurrence of ordovician carbonates in wells offshore labrador (bell & howie 1990) and onshore at ‘fossilik’ c. 65 km east of maniitsoq in west greenland (stouge & peel 1979; [9] on the 1:2 500 000 map). the crust under davis stræde is therefore interpreted by the survey as being formed of thinned continental crust, in accordance with the interpretation of the distribution of crustal types farther south. 71 baffin bugt (69–77°n) the distribution of crustal types underlying baffin bugt has not yet been mapped out with certainty, not least because linear sea-floor spreading magnetic anomalies are not distinct in this region. in the central, deep part of baffin bugt refraction seismic experiments have shown that the crust is very thin, the m (moho) discontinuity lying only 11 km below sea-level. the thickness of the cover of sediments exceeds 4 km in all but the southernmost part of this region. seismic velocities in the crust below these sediments are in the range 5.7–7.0 km/sec (srivastava et al. 1981; balkwill et al. 1990), in agreement with those known from oceanic layers 2 and 3. gravity and magnetic evidence (see next paragraphs) is also consistent with the interpretation of the central part of baffin bugt as being underlain by oceanic crust (e.g. balkwill et al. 1990). however, geophysical evidence has been presented which indicates that in north-western baffin bugt continental crust is replaced oceanwards by a layer of serpentinised mantle, which would account for the lack of distinct linear magnetic anomalies in this area (reid & jackson 1997). the weakness of linear sea-floor spreading magnetic anomalies could also be attributed to very oblique spreading (see roots & srivastava 1984) and to the dampening effect of the thick sedimentary cover. in spite of the weakness of the magnetic anomalies, oakey (2005) has succeeded in interpreting linear anomalies between 69° and 73°n and divided these into two directions, nnw–sse and nw–se, corresponding respectively to the paleocene and eocene anomalies in the labrador sea. oakey’s interpretation has been accep ted in the recently published 1:5 000 000 map of the entire polar region north of 60°n (harrison et al. 2008), on which the oceanic crust in baffin bugt is divided into crust of paleocene and eocene age. the same interpretation is shown in fig. 49a. the extinct spreading axis located by chalmers & pulvertaft (2001) is the axis of eocene spreading. oakey (2005) agrees with chalmers & pulvertaft (2001) that the major transfer faults along which the spreading axis was displaced, trend c. n–s. in the absence of easily recognisable sea-floor spreading magnetic anomalies, the landward limit of proven oceanic crust cannot yet be placed with any degree of confidence. existing released geophysical data are not sufficient to allow any interpretation of the position, width and nature of the continent–ocean transitional zone. on the 1:2 500 000 map the area underlain by continental crust was delineated on the evidence of crustal thickness and structural style (large extensional faults and rotated fault blocks), the latter being known from the kanumas reconnaissance seismic survey (whittaker et al. 1997). in southern baffin bugt a recently acquired nw–se refraction seismic line has revealed a material with high velocity (6.8 km/s) underlying sediments at c. 68°40΄n, east of the n–s transform fault running n–s at c. 60°w (t. funck, personal communication 2009). this could be oceanic crust, but at present neither the age nor the north–south extent of this high velocity layer is known. nares stræde (77–82°n) the nature of the geological structure underlying nares stræde, the linear seaway separating greenland from ellesmere island (fig. 49b), has for some time been a controversial subject (dawes & kerr 1982; tessensohn et al. 2006). geophysicists have argued that the strait is the site of a major transform fault with a left-lateral displace ment of more than 100 km that accommodated the opening of the labrador sea and baffin bugt during the palaeogene, and also of movement normal to the strait (e.g. srivastava 1985). in contrast, geologists familiar with the surrounding onshore geology find that there is no significant lateral displacement in smith sund (smith sound). several geological markers correlate perfectly from the thule-inglefield land area in greenland to south-east ellesmere island south of the bache peninsula (dawes 2009a). however, in the kennedy kanal (kanal = channel) and robeson kanal and along the north-west coast of kennedy kanal, both thrusting and sinistral strike-slip faulting have been observed (tessensohn et al. 2006), and a strike-slip displacement of up to 70 km is considered possible (harrison 2006). the apparent contradiction between what can be interpreted in kennedy kanal and what is observed across smith sund can be resolved if south-east ellesmere island is geologically speaking part of the greenland plate. assuming this is the case, the tectonic junction between greenland and ellesmere island runs along the north-west side of kennedy kanal as far south as 80°n and then turns inland just north of bache peninsula, where the strikeslip displacement in and along the north-west coast of kennedy kanal is transferred into thrusting along the southern front of the eurekan orogen (fig. 49a). the cumulative shortening across the major thrust zones of the eurekan orogen has been estimated to be of the order of 100 km (de paor et al. 1989). recent work in baffin bugt has diminished the likelihood that nares stræde is the site of a major transform fault, because 1) the direction of transform fracture zones and hence spreading movement in this area is almost 72 north–south, i.e. at an angle of about 40° to nares stræde (see fig. 49a, b; wheeler et al. 1996; whittaker et al. 1997; chalmers & pulvertaft 2001; harrison et al. 2008) and 2) the width of the area of oceanic crust in baffin bugt may be much less than previously supposed (oakey 2005; harrison et al. 2008)). it may well be that baffin bugt spreading has been accommodated to a substantial degree by a series of rifts in the canadian arctic islands together with compression in the eurekan orogen, just as the gakkel ridge spreading axis terminates in a system of rifts and extension zones in the laptev shelf north of siberia (drachev 2000). the continental margin off north greenland, west of the morris jesup rise ice conditions in lincoln hav (the lincoln sea) off north greenland are the most severe anywhere in greenland waters. consequently, the only indications of what might underlie the sea were for many years provided entirely by airborne geophysical data. summarising the results of earlier work in the region, dawes (1990) concluded that while gravity data from lincoln hav suggest that the north greenland margin is underlain by thinned continental crust, the crustal structure of the offshore region is conjectural. for this reason a wide area is shown without colour on the 1:2 500 000 map. in 2006 geophysicists succeeded in acquiring two wide-angle refraction/reflection seismic lines on the sea ice north of ellesmere island and greenland (dahl-jensen et al. 2006). these seismic sections show that thinned continental crust continues northwards from the inner greenland shelf under a 2000 m deep channel to the southern end of the lomonosov ridge. consequently it is now believed that the entire area shown without colour as ‘type of crust unknown’ on the 1:2 500 000 map is underlain by thinned continental crust. the >2000 m deep area shown as underlain by oceanic crust older than 45 ma [f] on the 1:2 500 000 map is also likely to be underlain by continental/transitional crust, an inference supported by the fact that linear sea-floor spreading magnetic anomalies identified to the north do not continue into this area (engen et al. 2008). the study by dahljensen et al. (2006) also showed that a previously suspected deep sedimentary basin underlying lincoln hav does in fact exist (see later). offshore sedimentary basins (fig. 56) reflection seismic surveys have shown that large sedimentary basins occur offshore east greenland between latitude intervals 67–72°n and 75–80°n. in the intervening area there are extensive palaeogene basalts below which thick sedimentary successions have been tentatively interpreted, but these cannot be resolved in existing seismic data. offshore west greenland there are rift basins with substantial thicknesses of sediment as far north as 76°n, and also smaller basins in southern nares stræde. as just mentioned, under lincoln hav a major sedimentary basin has recently been identified (dahljensen et al. 2006; k. sørensen, personal communication 2009). north-east greenland shelf (72–80°n) the existence of thick sedimentary successions on the north-east greenland shelf was first suggested on the basis of interpretation of aeromagnetic data (thorning et al. 1982; h.c. larsen 1984). seismic and gravity data ac quired as part of the kanumas project confirmed the existence of major sedimentary basins under this shelf, but at the time the 1:2 500 000 map was printed, very few details had been released. later a summary of the structure and succession in the basins in this shelf has been published (hamann et al. 2005). the sediments on the shelf were deposited in two basins separated by a basement high, the danmarkshavn ridge (fig. 50a). judging from the known geology of the barents sea, the norwegian shelf and onshore northeast greenland, the age of these sediments is likely to be devonian to recent, with unconformities in the middle permian and at the base of the paleocene. in the inner basin the maximum thickness of the basin fill is c. 13 km, and the succession is thought to span the entire period between devonian and neogene. a profound unconfor mity separates the ?devonian–cretaceous section from the overlying paleocene and younger units (fig. 50b). the presence of a thick salt layer in the deep part of the basin is clearly shown by gravity lows that coincide with diapiric structures seen in the seismic sections (hamann et al. 2005). this salt formation passes into time-equivalent carbonates deposited on the platform to the west. by comparison to the nordkapp basin, norwegian bar ents sea, the salt is inferred to be of late carboniferous – earliest permian age (hamann et al. 2005). a very thick succession also occurs in the outer basin. here, however, it has not been possible to interpret layers lower 73 than ?jurassic–cretaceous, although there are indirect indications that older sediments occur here, just as they do in the inner basin. on the shelf between latitudes 72°15´ and 75°30´n extensive volcanic rocks of presumed palaeogene age have been interpreted from the aeromagnetic and seismic data (h.c. larsen 1990). in the near-shore area these are exposed at the seabed; eastwards they become increasingly deeply buried under younger sediments. it is considered almost certain that the pre-paleocene sediments interpreted to the north and south of this area continue beneath the volcanic rocks. high amplitude magnetic anomalies suggest the presence of igneous intrusions in the sedimentary and volcanic rocks just north of 72°n. 74 ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ greenland fracture zone ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ w andel sea basin d an m ar ks ha vn b as in th et is b as in vo lc an ic p ro vi nc e a b k ol de w ey p la tfo rm o ll b il lb kcl 80°n 78°n 76°n 74°n 72°n 70°n 250 km inland ice ss younger sedimentary basins (55–0 million years old) oceanic basalts (55–0 million years old) palaeogene basalts and intrusions onshore/offshore platform sediment palaeozoic–mesozoic onshore/offshore basin: cretaceous–cenozoic offshore wandel sea basin sediments onshore/offshore basin: upper palaeozic – mesozoic onshore/offshore caledonian and older rocks onshore/offshore basin with salt diapirs oceanic crust continental crust fold belts and foreland illb: inner liverpool land basin ollb: outer liverpool land basin fracture zone fault thrust fault▲ fault in the subsurface jan mayenfracture zone ja m es on l an d b as in a 0 2 4 6 8 10 12 plio-pleistocene neogene palaeogene cretaceous intrusion upper permian – jurassic devonian – middle permian caledonian basement ?lower palaeozoic 50 km koldewey platform danmarkshavn basin thetis basindanmarkshavn ridge d ep th in k m a b b fig. 50. a: map showing the sedimentary basins offshore north-east greenland, between scoresby sund (ss, 70°n) and kronprins christian land (kcl, 82°n). modified from hamann et al. (2005). b: cross-section of the north-east greenland shelf at c. 77°n (a–b in a). from hamann et al. (2005). liverpool land basin, central east greenland (69°30´–72°n) a very thick succession of sediments can be recognised in the seismic data offshore liverpool land. the sediments are particularly thick within the part of the area underlain by continental crust, where the base of the sediments cannot be identified on the existing data. the upper part of the sedimentary succession is a virtually complete cenozoic succession up to 6 km thick; this formed a large prograding wedge that spread out across both continental and oceanic crust from the mouth of scoresby sund. in the part of the area underlain by continental crust the cenozoic succession lies with angular unconformity on block-faulted and tilted sediments of pre-paleo cene (late palaeozoic – mesozoic) age (fig. 51). where the cenozoic sediments have prograded into the area underlain by oceanic crust, they are underlain by subaerial lavas seen as seaward-dipping reflectors in the seismic data (larsen & jakobsdóttir 1988; h.c. larsen 1990). blosseville kyst basin, east greenland (67–69°30´n) more than 4 km of post-middle eocene sediments occupy an elongate, coast-parallel sedimentary basin off the blosseville kyst. the sediments lie entirely on palaeogene basalts. in the area underlain by continental crust there are almost certainly mesozoic and paleocene sediments beneath the basalts, as there are onshore and farther to the south (ocean drilling program well 917a; h.c. larsen et al. 1994a). however, it is not possible to interpret the geology underlying the basalts on the basis of existing seismic data. southern west greenland (60–68°n) several large, more or less coast-parallel, rift basins occur offshore west greenland between c. 62° and 68°n. smaller basins south of 62°n have yet to be mapped properly. the earliest sediments that can be confidently interpreted in these basins are preand syn-rift sequences up to 3 km or more in thickness, the kitsissut and appat sequences (fig. 52); by analogy with the labrador shelf, these are believed to be early cretaceous (barre mian–albian) in age (chalmers et al. 1993; chalmers & pulvertaft 2001). the abundance of ordovician carbonates in dredge samples collected on the davis strait high (dalhoff et al. 2006) indicates that ordovician limestones overlie basement locally in the region. as noted above, this possibility is supported by the occurrence of ordovician carbonates in wells offshore labrador (bell & howie 1990) and onshore at ‘fossilik’ c. 65 km east of maniitsoq in west greenland (stouge & peel 1979; [9] on the 1:2 500 000 map). 75 pleistocene late miocene _ pliocene late miocene oceanic crust continental crust (including palaeozoic– mesozoic sediments) crystalline rocks early _ middle miocene late oligocene _ early miocene eocene _ early oligocene 5 4 3 2 1 approximate position of ocean-to-continent transition inner liverpool land basin outer liverpool land basin twt sec sediments ew 20 km middle _ late miocene east greenland escarpment rift unconformity eocene ? pseudoescarpment fig. 51. e–w cross-section across the transition from continental to oceanic crust and the overlying liverpool basin at c. 71°n, east greenland. from larsen (1990 plate 6, profile e). the appat sequence is overlain by a widespread upper cretaceous mudstone-dominated succession, the kangeq sequence (fig. 52), the upper part of which was penetrated in the ikermiut-1 well (66°56´n, 56°35´w) and in the qulleq-1 well (63°49´n, 57°27´w). a major hiatus spanning the interval campanian – early paleocene (nøhr-hansen 2003) probably reflects the same episode(s) of faulting, uplift and erosion as are recorded in the succession in the nuussuaq basin to the north (see pp. 57–60). following these disturbances, fan sands intercalated with mudstones were deposited. early palaeogene volcanism gave rise to local ‘shields’ – the maniitsoq and hecla highs. deposition of mudstones continued into the early eocene, but from the middle eocene sedimentation was dominated by coarser clastic sediments deposited mainly in southwards prograding sequences. a major middle eocene (early lutetian) hiatus occurs in the three northernmost west greenland wells (hellefisk-1, ikermiut-1 and kangâmiut-1; nøhr-hansen 2003), and another major hiatus spans at least the oligocene (a.b. sørensen 2006). during the eocene, compressional structures developed in the area west of the ikermiut-1 well as a consequence of transpression along the transform ungava fracture zone, west of the 1:2 500 000 map boundary (chalmers et al. 1993). a study of dated minor intrusions onshore west greenland has shown that intrusion forms and melt compositions changed with time, depending on increasing lithospheric attenuation (l.m. larsen et al. 2009). significantly, this study suggests that by late jurassic – earliest cretaceous time thinning of the lithosphere had reached a stage when sedimentary basins could begin to form in the nascent labrador sea. central west greenland (68–73°n) the palaeogene basalts exposed onshore in the disko – nuussuaq – svartenhuk halvø area continue offshore where they have been mapped from seismic and magnetic data over the entire region between latitudes 68° and 73°n (gregersen & bidstrup 2008). in the eastern part of this region the basalts outcrop at the seabed and have been sampled by dredging, but to the west they become increasingly buried under a cover of eocene and younger sediments. while the upper surface of the basalts can usually be mapped easily from the seismic data, the base of the basalts is difficult to interpret and it is uncertain just where the basalts finally thin out and disappear to the west. nevertheless, from newer seismic data acquired by the oil industry it can be seen that a substantial thickness of mesozoic sediments underlies the basalts (gre gersen & bidstrup 2008). north-west greenland (73–77°n) north of 73°n seismic data acquired as part of the kanumas project confirmed the existence of a very deep graben or half-graben in the west and south-west part of melville bugt (melville bay; fig. 53; whittaker et al. 1997). this had earlier been outlined from aeromagnetic and gravity data acquired in the late 1960s and 76 mid-eocene – recent sediments lower eocene sediments paleocene basalts kangeq sediments appat/kitsissut sediments and sills basement (and possible older sediments) 1 0 2 3 4 ap pr ox im at e de pt h in k m 5 6 hecla high fylla fault complexnuuk basinlady franklin basin 25 km sea level lower eocene basalts paleocene sediments ew fig. 52. zig-zag cross-section through the area west of nuuk at about 64°n based on interpreted and depth-converted seismic data from an unpublished interpretation by j.a. chalmers. early 1970s. the more recent data have also revealed several other graben and half-graben structures extending to the northern limit of the survey at 76°30´n. in the melville bugt graben the thickness of sediments exceeds 12 km. by analogy with the onshore geology of west greenland and north-east canada (bylot island), the main phase of rifting is thought to have taken place in the cretaceous, prior to sea-floor spreading in baffin bugt. later, northern parts of the area were subjected to marked inversion. small sedimentary basins with up to 4 km of ?creta ceous–neogene sediments occur in an area known as north water bay, between kitsissut (carey islands) and smith sund (neben et al. 2006). the two easternmost of these basins strike nw–se; these are on line with extensional basins mapped farther south in northern melville bugt (whittaker et al. 1997) and hence are not likely to belong to any pull-apart system (pace neben et al. 2006). sediments belonging to the thule supergroup underlie the younger sediments and appear to be continuous right across north water bay (funck et al. 2006). offshore north greenland; the lincoln sea basin as already mentioned, ice conditions in lincoln hav are the most severe anywhere offshore greenland, and until recently airborne geophysical data provided the only hints that a major sedimentary basin lies underneath this sea (sobczak & stephens 1974; kovacs 1982; mcmillan 1982). recently the two wide-angle refraction/reflection seismic traverses successfully carried out by dahl-jensen et al. (2006) not only established the continental nature of the crust between the greenland margin and the southern end of the lomonosov ridge but also revealed an up to 14 km deep sedimentary basin consisting of two layers interpreted to be part of the arctic continental terrace wedge, under which there is a 9 km thick layer interpreted to be the offshore continuation of the upper palaeozoic – mesozoic succession in the sverdrup basin in the canadian arctic islands. 77 0 2 4 6 8 sw ne twt sec ? ? ? ? ? baffin bugt ? oceanic crust/ transition zone kivioq ridge kivioq basin melville bugt ridge melville bugt graben melville bugt fault glacial deposits thermal subsidence post-drift early syn-rift break-up/syn-drift basementlate syn-rift sediments crystalline rocks 25 km fig. 53. representative cross-section offshore north-west greenland at c. 75°n compiled from seismic reflection data. from whittaker et al. (1997). mineral raw materials in greenland occur in a series of different geological environments that include sedimentary deposits, metamorphic crystalline rocks, and volcanic and plutonic rocks. an overview of selected mineral occurrences with their place names are shown in fig. 54. mining activities have been carried out in greenland since the middle of the 19th century, with the cryolite mine at ivittuut as the only long-term mine; it was in operation for a period of 130 years. the cryolite deposit was associated with a granite intrusion in the meso proterozoic gardar province of south greenland (p. 38), and represents an example of a very rare type of mineralisation, of which there are only very few similar deposits in the world (pauly & bailey 1999). other mining activities in greenland have exploited more common types of mineralisation. the two most important ones were both lead-zinc deposits – one at mestersvig in east greenland was associated with quartz veins of probable palaeo gene age, and the other at maarmorilik in central west greenland was a stratabound mineralisation in the palaeoproterozoic mârmorilik formation (p. 27). most recently a gold occurrence in kirkespirdalen in the palaeoproterozoic ketilidian orogen, south greenland, has been mined (nalunaq gold mine, see below). mining activities have so far been very limited in green land considering the expected potential of such a large area. however, systematic exploration did not commence until the late 1950s when new legislation governing the mineral sector was introduced to encourage the mining industry to undertake exploration. this was intensified with the introduction of home rule status for greenland in 1979. in recent years exploration activities have concentrated on prospecting for gold, base metals, platinum elements, molybdenum, iron ore and diamonds. gold exploration has focused on the archaean and palaeo proterozoic precambrian shield of west greenland and the palaeogene skaergaard layered gabbro intrusion (p. 62) in southern east greenland (andersen et al. 1998; secher et al. 2007). a major new gold province in the palaeoproterozoic ketilidian orogen forming the southern tip of greenland (steenfelt 2000; stendal & frei 2000; stendal & secher 2002) was first recognised by panning of stream sediments, and in 1992 visible gold was found in quartz veins transecting mafic supracrustal rocks. the nalunaq gold mine in kirkespirdalen (a valley north-east of nanortalik) operated from 2004 to 2008, but the mine is presently (2009) placed on ‘care and maintenance’. another find substantiating the interpretation of the ketilidian orogen as a gold province was made in southernmost south-east greenland, where gold mineralisation was found in a quartz-bearing shear zone cutting a sequence of mafic to andesitic extrusives and intrusives and associated sedimentary rocks. the promising gold mineralisations in south greenland are situated at the southern border of the julianehåb batholith (p. 29), which is also the root zone of a former volcanic arc (chadwick & garde 1996; garde et al. 1998, 2002; mccaffrey et al. 2004). exploration for base metals in recent years has focused on showings in the lower palaeozoic franklinian basin and ellesmerian fold belt of north greenland. a massive sulphide deposit with lead and zinc was discovered in 1993 at citronen fjord in peary land (fig. 54; van der stijl & mosher 1998). it occurs as stratiform sheets in a folded sequence of dark argillaceous rocks of the upper ordovician to lower silurian amundsen land group. the citronen fjord deposit is located at the eastern end of the franklinian basin, which extends across north greenland into arctic canada and is known to be a significant prospective zone which includes the polaris zinc-lead mine in canada. diamond exploration has focused on the archaean and palaeoproterozoic crystalline shield areas of west green land. the kimberlite province in this region in cludes the archaean craton and areas of archaean rocks farther north reworked during the palaeoproterozoic. the province contains various meso–neoproterozoic ultramafic lam prophyres (uml) as well as mesozoic uml-intrusions. 78 mineral deposits facing page fig. 54. map of known mineral occurrences in greenland. the presentation combines the geological map with the knowledge of mineral prospects and deposits gathered over more than 150 years. the main geological divisions on the geological map cover periods from archaean to palaeogene and are summarised in the legend. from geus department of economic geology. 79 e l l e s m e r e i s l a n d c a n a d ac a n a d a 250 km nalunaq au niaqornaarsuk auamitsoq graphite ilímaussaq zr, be, ree, nb, u, th ivittuut cryolite taartoq au kobberminebugt (‘josva’) cu grønnedal-ika fe motzfeldt sø ta, nb sinarsuk v, ti kangerluluk au illorsuit u stendalen cu, ni, ti fiskenæsset cr, pt, ruby isukasia fe isua au ivisaartoq w majuagaa diamond seqi olivine amikoq pt, pd, os qaqqaarsuk ree, nb, p sillisissanguit ni, pt attu au sarfartoq nb, ta storø au qussuk au eqi au hammer dal fe, pt maarmorilik (‘black angel’) zn, pb, ag karrat au, cu, zn moriusaq ti washington land zn, pb, ag inglefield land fe, au,cu ymer ø sb, w, au clavering ø pb, zn, fe brogetdal cu citronen fjord zn, pb qullissat coal nuussuaq coal arveprinsen ejland cu, zn ilukunnguaq ni, pt saqqaq au itilliarsuk, itilli fe, au, cu, co, ni eqalussuit graphite melville bugt fe langø graphite navarana fjord zn, barite naternaq cu, zn kangerlussuaq diamond mestersvig (‘blyklippen’) pb, zn malmbjerg mo, w bredehorn barite milne land zr, ree, ti devon dal cu flammefjeld mo, au, ag kap edvard holm au, pt tasiilaq ni, cu, pt skærgaard au, pd, ti, v karstryggen celestite i n l a n d i c e ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ▲ ▲ ▲ ▲ ▲ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ■ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ■ ■ ■ ■ ■ ■ ■ ■ e l l e s m e r e i s l a n d c a n a d a carboniferous–cretaceous sediments, north-east greenland palaeogene basalts cretaceous–palaeogene sediments, nuussuaq basin in west greenland and kangerlussuaq basin in east greenland carboniferous–palaeogene sediments, wandel sea basin in eastern north greenland carboniferous–cretaceous sediments, jameson land basin in east greenland devonian basin of north-east greenland shelf trough caledonian orogenic belt palaeoto neoproterozoic sediments and volcanic rocks palaeoproterozoic orogenic belts archaean craton intrusive complexes: palaeogene in east greenland, mesoproterozoic in south greenland (gardar province) fault, thrust mine site abandoned mine lower palaeozoic sediments, north greenland, franklinian basin i c e l a n d ▲ ▲ ▲ ▲ a large number of these intrusions contain diamonds, and intensive exploration activity has been in progress since the 1990s (secher & jensen 2004; jensen et al. 2004; nielsen et al. 2009). at present more than 1000 occurrences of diamondiferous kimberlite dykes have been found, and the largest diamond so far discovered is of 2.5 carat. the archaean block also contains three large intrusive carbonatite complexes with a resource potential for various speciality commodities such as niobium and tantalum. mineral occurrences in specific geological settings significant occurrences of a broad range of metallic and industrial minerals are present in all the principal geological provinces in greenland, ranging in age from archaean to quaternary (fig. 54; schønwandt & dawes 1993; stensgaard & thorning 2009). in broad terms these can be related to five main settings: archaean–palaeoproterozoic high-grade regions mesoproterozoic intracratonic intrusions palaeozoic orogenic belts upper palaeozoic – mesozoic basins late phanerozoic intrusions. the following description covers the principal active and former mines and some significant prospects (fig. 54); at present (2009) there are two active mines, although one is ‘on hold’. on the printed map sheet the locations of only four abandoned mines are shown (ivittuut cryolite mine in south-west greenland, maarmorilik zinc and lead mine in central west greenland, mestersvig lead mine in central east greenland and qullissat coal mine in central west greenland). promising prospects are a large zn-pb mineral occurrence in citronen fjord in north greenland, a banded iron occurrence at isukasia in the nuuk region, a zirconium and rare-earth element prospect/deposit in the ilímaussaq complex in south greenland and the malmbjerg molybdenum deposit in central east greenland. information about a large number of mineralised localities is available in the continually updated greenland mineralisation data bank at the survey (see: www. geus.dk/gmom: greenland mineral occurrence map on-line). archaean–palaeoproterozoic high-grade regions at isukasia (isua supracrustal sequence [69]; fig. 54) north-east of nuuk, a major archaean banded iron formation is composed of interlayered magnetite and chert (fig. 5). the deposit, which is partly covered by the in land ice, has been drill tested and a minimum tonnage of 1900 million tonnes grading 32.9% fe is estimated (niel sen 1976; appel 1991). the nuuk region in southern west greenland has revealed a good potential for gold mineralisations (stensgaard & stendal 2007). the gold occurs in the supracrustal parts of the archaean craton, which largely consists of amalgamated islands arcs, which gradually merged into micro-continental blocks (windley & garde 2009). the supracrustal belts reflect both island-arc and ocean-floor environments, and also contain ultramafic to mafic magmatic intrusions. gold showings occur in a nne-trending belt along the fjord nuup kangerlua (godthåbsfjord) from nuuk to isukasia, and a multidisciplinary approach has now resulted in division of the occurrences into three main groups (stensgaard & stendal 2007). the occurrences contain up to 3–7 g/t au with local grades of up to 20 g/t au. intensive exploration with drilling both north and south of the fjord has been undertaken since 2003. a series of gold occurrences has also been found in the crystalline basement rocks of the nagssugtoqidian orogen north-east of disko bugt, central west greenland. here archaean orthogneisses with their archaean and palaeoproterozoic cover rocks have been variably affected by the c. 2.0 to 1.75 ga nagssugtoqidian orogeny. the gold is hosted in archaean metasedimentary and meta volcanic rocks and may be either stratabound or located in veins, breccias or shear zones (steenfelt et al. 2004). the gold values are modest with only a few ppm in the mineralised zones, but there is a potential for further occurrences in the investigated region. a folded and metamorphosed archaean anorthosite complex [85] at qeqertarsuatsiaat/fiskenæsset, southern west greenland (the fiskenæsset complex; fig. 54) hosts widespread chromite-bearing layers (ghisler 1976). the complex, which has a strike length of more than 200 km and an average thickness of 400 m, has an estimated po tential of 100 million tonnes of low-grade chromium ore. enhanced precious metal values have been reported from the ultramafic parts of the complex (appel 1992). ruby-bearing rocks occur at several localities in the fiskenæsset complex, where the corundum/ruby occurs in zones close to the contact between anorthosite and 80 amphibolite/ ultrabasite. a mining company has obtained an exclusive exploration licence for rubies, and tests have shown that some types of ruby and sapphire are of gem quality (secher & appel 2007). a neoarchaean iron province occurs in the coastal areas along melville bugt in north-west greenland. geographically it is the largest iron province in greenland (dawes 2006) and is traceable in a wnw–ese-trending belt for c. 350 km. the belt contains magnetite and hematite in quartz-banded iron formation (bif), massive lenses and layers with iron oxides and disseminated iron minerals in schists. bif occurs in units of varying thickness, ranging from less than a metre to 40 m where iron concentrations typically are 30–35%. at maarmorilik, central west greenland, the black angel (fig. 54) zinc-lead ore bodies hosted in palaeo proterozoic marble (p. 27; lower part of [62]) were mined in the period 1973–1990. production totalled c. 11 million tonnes ore grading 4.0% pb, 12.6% zn and 29 ppm ag. the deposits were almost exhausted in that period, but a re-establishment of the mine is planned for 2010–2011, based on the remaining ore in pillars in the mine combined with other marble-hosted lead-zinc prospects in the area (thomassen 1991, 2006; minex 2008b). prior to the pb-zn mining, some 8000 tonnes of marble were quarried at maarmorilik in 1936–1971. the nalunaq gold mine in kirkespirdalen, north-east of nanortalik, south greenland (figs 54, 55) is a small, high-grade gold deposit associated with up to 1.8 m wide quartz-veins in a major shear zone. the deposit is an orogenic-type gold mineralisation (mesothermal lode gold) hosted in palaeoproterozoic amphibolite facies volcanic rocks within the ketilidian orogen. the mine was opened in 2004, but mining was placed on ‘care and maintenance’ by the owner at the end of 2008. in 2009 a new company is negotiating to take over the 81 level 450 level 400 level 350 level 300 mining camp exposed quartz vein fig. 55. view of the mining area around the nalunaq gold mine in south greenland. access to the underground mine is at four levels of 300, 350, 400 and 450 m. the mine was opened in 2004. photo: bureau of minerals and petroleum, nuuk. mine and start production again. production in 2006 totalled 108 000 tonnes of ore with an average gold grade of 17.9 grams per tonne (secher et al. 2008). several large, homogeneous, olivine-rich (dunitic) bodies occur in the archaean gneiss terrain some 90 km north of nuuk. an opencast mine, the seqi olivine mine, was opened in 2005 in the northernmost part of the niaqunngunaq/fiskefjord fjord system (see fig. 54).this homogeneous deposit contains at least 100 million tonnes of high-quality olivine. mesoproterozoic intracratonic intrusions cryolite hosted in a gardar granite stock (part of [56]) at ivittuut, south-west greenland (fig. 54) was worked from 1858 until 1987, and a total of 3.7 million tonnes ore grading 58% cryolite was quarried from an open pit. in addition to cryolite, galena, chalcopyrite and siderite were extracted as by-products. the main ore body is now exhausted, but there are indications of deep-seated reserves in the area (bondam 1991). in the gardar ilímaussaq alkaline intrusion (part of [56]; see fig. 19; fig. 54) east of narsaq in south green land, a deposit with rare metals such as niobium, tantalum, zirconium, yttrium, rare-earth elements, lithium and beryllium and accessory uranium and thorium has been delimited by diamond drilling, indicating a reserve of 56 million tonnes of u with a grade of 365 ppm (nyegaard 1979). the rare-earth and other special elements are concentrated in the final, highly volatile, products of the magmatic differentiation. the last rocks to solidify are therefore relatively rich in elements such as niobium, tantalum, zirconium and rare-earth elements (bondam 1995). the motzfeldt centre east of narsarsuaq is another of the large intrusive syenite complexes in the gardar province. this centre has preserved accumulations of the mineral pyrochlore which has a high content of tantalum and niobium. the deposit has been explored and is rated as a ‘low-grade – large tonnage’ type of re source, with 600 million tonnes of ore with 120 g/t tantalum and 130 million tonnes of ore with 1400 g/t niobium. the deposit is believed to be one of the largest tantalum deposits in the world (h. sørensen et al. 2006b). late neoproterozoic (604–555 ma old) in situ kimberlite intrusions and ultramafic lamprophyre dykes are widespread as dykes and sheets in the region around the fjord kangerlussuaq / søndre strømfjord, west greenland (secher et al. 2009). they are found on both sides of the boundary between the archaean craton to the south and the palaeoproterozoic nagssugtoqidian orogen to the north. the intrusions have been known since the mid1960s and the first microdiamonds were found in stream sediments in the sarfartoq river. the finds of micro diamonds led to intensive prospecting, and at present several microdiamond sites have been reported from outcropping kimberlitic rocks (minex 2002, 2007a, 2008a, b; secher & jensen 2004). diamonds and diamond indicator minerals have since been found in numerous samples of stream sediments, in boulders and in situ kimberlites, both in the sarfartoq area and in the maniitsoq/sukkertoppen area (65°n). in the sarfartoq area exploration has resulted in finds of commercialsized diamonds of up to 2.5 carats, and in 2008 the exploration company reported the recovery of another large diamond of c. 4 carat from a dyke. in the maniitsoq area recent field work has resulted in recognition of a large kimberlite dyke system with a combined length of more than 10 km. the known carbonatite intrusions of sarfartoq, qaqarssuk and tikiusaaq (legend numbers [59] and [61] on the 1: 2 500 000 map) within the alkaline province south of kangerlussuaq / søndre strømfjord have a recognised potential for accumulations of niobium and rareearth elements (secher & jensen 2004; kolb & stens gaard 2009). tikiusaaq was found in 2005 (steenfelt et al. 2006) and is thus not indicated on the map. palaeozoic orogenic belts south of citronen fjord (fig. 54) in peary land, north greenland, a large sulphide-rich zone hosts a major leadzinc-bearing, sedex-type, massive sulphide deposit in ordovician black shales (part of [24]) (kragh et al. 1997). diamond drilling since the discovery in 1993 and up to 2008 has yielded 44 km of core and has indicated a resource of more than 102 million tonnes grading 4.7% zn + pb at a 2% zn cut-off grade (van der stijl & mosher 1998; minex 2007b, 2008b). the deposit is located north of a prominent palaeo-escarpment separating carbonate shelf sedimentary rocks to the south from deepwater trough sedimentary rocks to the north. the prospect is currently (2009) under evaluation for opening of a mine in the near future. another discovery of zinc-lead-silver mineralisation in the same province has been made in washington land, western north greenland, where a galena occurrence is hosted in evaporitic lower ordovician carbonates in the platform succession (jensen 1998). 82 upper palaeozoic – mesozoic basins at qullissat (fig. 54) on disko, central west greenland, cretaceous sub-bituminous coal was mined during the period 1924–1972. a total of about 570 000 tonnes of coal was shipped before the mine was closed due to the low coal quality (schiener 1976). on nearby nuussuaq, more than 180 million tonnes of sub-bituminous coal distributed in layers more than 0.8 m thick have been indicated by surface investigations and limited drilling (shekhar et al. 1982). late phanerozoic intrusions the 54.5 ma old skaergaard layered gabbro intrusion (fig. 54) at kangerlussuaq, north-east of ammassalik in southern east greenland, hosts a major deposit of lowgrade palladium, platinum and gold (bird et al. 1991). intensive diamond drilling totalling 42 drill holes with a combined length of more than 21 km, has shown a resource of more than 1520 million tonnes grading 0.21 ppm au, 0.61 ppm pd and 0.04 ppm pt (thomassen & nielsen 2006; secher et al. 2007). the mineralisation is hosted in a 100 m thick zone with gold and platinumgroup elements accumulated in five intervals with thicknesses of several metres. in these intervals concentrations of gold and platinum-group elements are much higher than the average figures given above. titanium, vanadium and iron are important additional commodities in the middle of the mineralised zone, and a test profile across the deposit indicates average contents of 6.6% tio2, 0.13% v2o5 and 19% fe2o3. similar mineralisation is known in other nearby intrusions. lead-zinc-bearing quartz veins, probably of palaeogene age, occur in lower permian sediments near mestersvig in east greenland (fig. 54). one of these occurrences, the blyklippen deposit, was mined in the period 1956–1962. after production of 560 000 tonnes ore grad ing 11.1% pb and 8.6% zn the deposit was exhausted (harpøth et al. 1986; thomassen 2005). a large porphyry-molybdenum deposit of miocene age occurs at malmbjerg (fig. 54) south of mestersvig, east greenland, hosted in an intrusive complex [53]. ore resource calculations were based on 22 km of diamond drill cores which indicate a tonnage of 150 million tonnes grading 0.23% mos2 and 0.02% wo3 (harpøth et al. 1986). a re-evaluation of the deposit was initiated in 2004 and has confirmed the earlier tonnage estimates. a mining company obtained an exploitation licence in 2009 and aims at opening an open pit mine in the near future with a production rate of c. 10 000 tonnes molybdenum per year (minex 2007b, 2008b). other less well-investigated porphyry-molybdenum occurrences exist in the east greenland palaeogene volcanic province (geyti & thomassen 1984). 83 the petroleum potential of greenland is confined to the sedimentary basins of phanerozoic age. onshore, such basins occur in north greenland, north-east and central east greenland, and central west greenland. off shore, large sedimentary basins are known to occur off both east and west greenland (fig. 56). no proven commercial reserves of oil or gas have been found to date (2009), but so far only seven exploration wells have been drilled, six offshore southern west greenland between latitudes 63°49´n and 68°n, and one onshore, on nuus suaq at 70°28´n in central west greenland (christiansen et al. 1997; pulvertaft 1997). in recent years there has been much interest in petroleum exploration mainly offshore west greenland, where a number of exploration licences (fig. 57) have been granted to consortia of both large and small oil companies (f.g. christiansen, personal communication 2009). in the coming years focus will also be directed towards baffin bugt off north-west greenland. recent investigations include acquisition of geophysical data (seismic and airborne magnetic/gravity surveys) and seabed sampling. another main target for future oil and gas exploration will be the shelf areas of north-east greenland, where geophysical investigations have revealed the existence of a number of large sedimentary basins (hamann et al. 2005). based on data from the adjacent onshore areas it may be assumed that source rocks, reservoirs and seals are likely to occur here, and that several play types are present. the geophysical investigations in north-east greenland both onshore and offshore have up to now only been carried out at reconnaissance level, and the area must still be characterised as essentially unexplored for oil and gas. the greatly increased interest in recent years for petroleum exploration in greenland has been supported by its relatively high ranking given by the arctic petroleum appraisal of the united states geological survey. here, north-east greenland was selected as a prototype for an evaluation of this and similar other areas in the circum-arctic region (f.g. christiansen, personal communication 2009). the undiscovered resource estimates for both oil and gas are quite high and the danmarkshavn basin (see p. 74) in particular is mentioned as a promising area (gautier et al. 2009). a brief summary of the petroleum-geological features of the main sedimentary basins is given on the follow pages. onshore basins franklinian basin, north greenland (80–83°n) the franklinian basin of north greenland (see p.45) is the eastern continuation of the cambrian–devonian franklinian basin of the canadian arctic islands. it consists of a belt of flat-lying, shallow-water carbonate rocks to the south and a northern belt of deep-water folded sedimentary rocks. good type ii (oil-prone) shaly source rocks are known in both lower–middle cambrian and lower silurian outer shelf terrigenous and carbonate mudstones. potential reservoirs include lower and middle cambrian shelf sandstones and lower silurian reef and platform margin carbonate build-ups (stemmerik et al. 1997). the rocks in the north are probably postmature due to the thermal influence of the ellesmerian orogeny (see p. 49), but oil has been preserved in the southern areas and can now be seen as asphalt residues in pores and fractures in various carbonate rocks. the most promising play involves long-distance migration up-dip from middle cambrian source rocks into lower cambrian shelf sandstones (christiansen 1989). late palaeozoic – mesozoic basins, eastern north greenland (80–83°n) deposits in the wandel sea basin comprise a succession of sedimentary rocks which were laid down along the northern and north-eastern margin of the greenland shield (see p. 54). the development spans a period from early carboniferous to palaeogene and includes three main phases of basin formation. the region is transected by a major nw–se fault zone dividing the area into two blocks with different structural, depositional and thermal histories and hydrocarbon potential (stemmerik 84 petroleum potential of greenland facing page fig. 56. simplified map showing the onshore and offshore sedimentary basins with petroleum potential of greenland. distribution of crustal types can be seen in fig. 48 a. ellesmere island (canada), iceland and svalbard are shown as light brown: undifferentiated. map based on henriksen et al. (2000), chalmers & pulvertaft (2001) and hamann et al. (2005). 85 62° ▲ ▲ ▲ ▲ ▲ ■ ▲ ▲ ▲ ▲ ▲ ▲ ▲ melville bugt melville bugt canadacanada eell llee ssmm eerree ii ss llaanndd nares s t ræ de melville bugt d e n m a r k s t r a i t labrador sea at lant i c ocean el le sm ere is land 50°60°70° 40°w 30° 20° 10° 0° 10°ø 10° 80° 0° 78° 76° 10° 74° 66° 70° 64° ice land illoqqortoormiut / scoresbysund 70° 74° 72° 68° 60 ° sisimiut 64° nuuk qaanaaq 50° 20° nares s t ra i t melville bugt disko bugt qaqortoq upernavik daneborg 40°w 35° reykjavik 60° 62° 70° 76° canada ilulissat alert 80° 78° 90° 72° 90° 80° m elville bay graben kivioq basin ikermiut basin ka ng aa m iu t ba sin n uu k ba sin paam iut basin lady franklin basin sisimiut basin liverpool land basin thetis basin d anm arkshavn basin wandel sea basin lincoln sea basin b o rd er w it h c an ad a aa sia at ba sin 250 km g r e e n l a n d s e a ages in million years basins with petroleum potential onshore areas younger sediments (400–0) cambro-silurian sediments (540-400) offshore basins sediments, large basins (400–0) basalts and intrusives (60–50) ellesmerian fold belt (350) caledonian fold belt (420) older sediments and magmatic rocks (1740–420) proterozoic basement (2000–1750) archaean basement (3800–2550) other basins and crystalline rocks onshore areas oceanic basalts (65–0) basalts and intrusives (60–30) younger fold belts and basement locally covered by sediment offshore areas station nord sva lbard davis strait 66° baffin bay ▲ ▲ ▲ ▲ 68° spreading axis inactive active extensional fault compressional fault thrust fault zone deep borehole bl os se vil le ba sin tasiilaq/ammassalik et al. 2000). to the north of the fault zone the prepaleocene sedimentary rocks are considered postmature with respect to petroleum generation and of limited economic interest. this is also likely to be the case on the nearby shelf in eastern north greenland where a similar sedimentary succession is expected to be present. in contrast, onshore sedimentary rocks in the coastal region south of the fault zone are early mature to immature and therefore there might be a prospective zone in the offshore region along the expected continuation of the fault zone at c. 80°n (fig. 37). carboniferous and permian reservoir rocks occur onshore, but source rocks have not been identified. late palaeozoic – mesozoic rift basins, north-east greenland (72–76°n) the main source rocks in these north-east greenland basins are: (1) upper carboniferous type i–ii (highly oil-prone – oil-prone) mudstones with very high genera tive potential but restricted lateral extent, (2) upper permian type ii marine mudstones with wide areal extent and high generative potential, and (3) upper jurassic (kimmeridgian) marine mudstones which are mainly gas-prone in onshore outcrops but are likely to be highly interesting oil-prone source rocks on the continental shelf to the east (hamann et al. 2005). reservoir lithologies include upper carboniferous fluvial sandstones, upper permian carbonates, upper jurassic sandstones, and uppermost jurassic – lower cre taceous syn-rift conglomerates and sandstones. the basins are partially fault bounded and tilted, and there are both stratigraphical and structural plays. from regional mapping and maturity considerations an area of about 6000 km2 is considered to have potential prospectivity (stemmerik et al. 1993), but at present there are no seismic data on which to base a more stringent evaluation. a more detailed understanding of the jurassic bio stratigraphy and depositional models with their potential for source and reservoir lithologies has been gained in recent years (ineson & surlyk 2003; stemmerik & stouge 2004). jameson land basin, central east greenland (70°30´–72°n) the jameson land basin, which extends over an area of about 10 000 km2, is covered by a 1798 km seismic survey, carried out by atlantic richfield company (arco) in 1985–89, and consequently is better known than the basins to the north. the structural history of the basin is also different in that rifting began in the devonian and ended in the mid-permian; late permian – mesozoic de position in the basin was governed by thermal subsidence. in addition to the source rock intervals known to the north (christiansen et al. 1992), an important lowermost jurassic lacustrine type i–ii source rock (highly oil-prone – oil-prone) occurs in jameson land (dam & christiansen 1990). potential reservoirs are upper carboniferous (and possibly older) fluvial sandstones, upper permian carbonates, and lower jurassic deltaic sandstones. apart from an upper carboniferous tilted fault block play, play types are stratigraphic. arco stopped their exploration activities at a time when the potential seemed restricted to a permian play in northwest jameson land. later reinterpretation by geus of the seismic data, supplemented by new field work, analyses of source and reservoir rocks and modelling also suggest a possible lower jurassic play in central jameson land (dam et al. 1995). the main risk factor in the jameson land basin is the effect of palaeogene and neogene uplift that amounts to 2 km or more (mathiesen et al. 1995). cretaceous–palaeogene basin, central west greenland (69–72°n) source rocks in outcrop are mainly gas-prone, but the discovery of surface oil showings in vesicular basalts over a large area extending from northern disko to southeast svartenhuk halvø, as well as the occurrence of oil in three of the five core holes drilled on western nuussuaq in 1993–1995, prove that source rocks capable of generating oil occur in this region. a 3 km deep wild-cat well (gro#3) was drilled in 1996 on south-western nuussuaq by the small canadian company grønarctic energy inc. the logs yield some indications of oil and gas, but without giving sufficient background for a continuation of the work (christiansen et al. 1999). in the region as a whole, biomarkers in the oils indicate that five types of oil are present, with source rocks of cretaceous–paleocene age (bojesen-koefoed et al. 1999). reservoirs in the area may be either cretaceous deltaic sandstones or uppermost cretaceous – lower paleocene turbiditic sandstones. 86 offshore basins north-east greenland shelf (75–80°n) an area of more than 125 000 km2 offshore north-east greenland is believed to have considerable petroleum potential. this view is based on extrapolation from the adjacent onshore area, where oil source rocks are present at several levels, and also from the northern north sea, west norwegian shelf and south-west barents sea, areas which were contiguous with the north-east greenland shelf before the opening of the green land–norwegian sea (tsikalas et al. 2005). the kanumas reconnaissance seismic survey of the shelf area confirmed that thick sedimentary basins occur on the shelf comprising possible devonian to neogene deposits with a thickness of up to c. 13 km (hamann et al. 2005). interpretation of the gravity and seismic data furthermore indicates that upper carboniferous – lower permian salt deposits are widespread between c. 77–79°n (stem merik & worsley 2005), as shown on the geological map and on fig. 50a. the east greenland succession on the shelf almost certainly includes upper jurassic and other source rocks. in the danmarkshavn basin (see fig. 50b) the jurassic sediments have been buried deeply enough to generate hydrocarbons. the succession is expected to include extensive, excellent quality source rocks, and trap structures include large-scale fault blocks (hamann et al. 2005). possible source rocks are correlatives of the following onshore occurrences: 1) organic-rich marine shales from the upper permian ravnefjeld formation considered to be good to excellent source rocks (christiansen et al. 1993), 2) marine jurassic shales of kimmeridgian age known as the hareelv formation in central east greenland (christiansen et al. 1992, 1993; surlyk 2003) and other equivalents to the world class upper jurassic source rocks known from the north atlantic region (christiansen et al. 1993; hamann et al. 2005), 3) upper triassic – lower jurassic lacustrine organic-rich shales from the kap stewart formation in central east greenland and 4) other source rocks may be found in lacustrine deposits of late palaeozoic age and from equivalents of middle jurassic coal deposits found onshore north-east green land. the arctic petroleum appraisal of the united states geological survey has rated the north-east greenland shelf region an area with major potential for oil and gas; most of the undiscovered resources are likely to be in the danmarkshavn basin (gautier et al. 2009). liverpool land basin, central east greenland (69°30´–72°n) up to 6 km of cenozoic sedimentary rocks unconformably overlie block-faulted upper palaeozoic – mesozoic sedimentary rocks in the inner (landward) part of the liverpool land basin. in the outer part of this basin oceanic crust occurs beneath a thick wedge of neogene and plio–pleistocene sedimentary rocks (h.c. larsen 1990; hamann et al. 2005). source rocks are likely to 87 sisimiut ilulissat nuuk nuussuaqdisko jameson land tasiilaq qaanaaq bor de r w ith ic ela nd greenland canada iceland 250 km existing licences blocks to be awarded kanumas preference area open area exploration well disko qaqortoq 76° 72° 80° 28°36° 44° 28° 44° 68° 64° 60° b or de r w ith c an ad a fig. 57. summary map of greenland showing the state of concession agreements for hydrocarbon exploration at the end of 2008 (awarded licence areas shown in red). licensing rounds were held in 2002, 2004, 2006 and 2007. licences in two smaller areas west of nuuk were awarded in 2003 and 2005, and a large region west of disko was awarded in seven block areas in 2007 to four groups of oil companies. the green areas are blocks to be awarded in connection with a licence call round in 2010. the violet areas are covered by a so-called ‘open-door’ policy where oil companies can apply for concessions at any time. the yellow areas show where a group of companies (the kanumas companies) have preferential positions, but where exploration is not expected to take place in the near future. from bureau of minerals and petroleum, nuuk. occur at several levels in the pre-cenozoic sedimentary rocks, but are probably postmature. nothing can be deduced about the nature of mudstones in the palaeogene. only a few weak structures have been observed in the cenozoic section, and the best traps are likely to be stratigraphic. blosseville kyst basin, east greenland (67–69°30´n) only the post-basalt cenozoic sedimentary rocks in the blosseville kyst basin are considered likely to have any potential for petroleum, since any sedimentary rocks underlying the basalts will be thermally postmature. the outermost sedimentary rocks overlie oceanic crust. trap structures occur where the sediments drape buried volcanic edifices, and it is likely that there are also stratigraphic traps. submarine fan sandstones fed from the land areas to the north and north-west are likely to be the best potential reservoirs in the area. source rocks are most likely to occur in the eocene – lower oligocene sedimentary rocks, which were deposited at a time when the area had only limited connections with the early atlantic ocean, a factor that would favour oxygen-deficient conditions (h.c. larsen 1985). west greenland southern west greenland was the first offshore area where companies were awarded exclusive licenses for hydrocarbon exploration. about 37 000 km of seismic data were acquired in the shallower parts of the area (water depths <500 m) in the early 1970s, and five wells were drilled. one well (kangâmiut-1, c. 66°n) encountered wet gas (chalmers et al. 1995), but the others were dry. with hindsight it can be seen that only the kang âmiut-1 well tested a viable structure (chalmers & pulvertaft 1993). in the 1990s exploration was resumed in the region, and more than 23 000 km of additional seismic data were acquired, extending knowledge of the geology into deeper water areas which appear to be the most prospective. a sixth well (qulleq-1, west of nuuk) was drilled by statoil in 2000; this yielded important new stratigraphic information but was dry (christiansen et al. 2001). in the last ten years, interest for oil and gas exploration in west greenland has been driven by the documentation of live petroleum systems onshore between 70°12´ and 71°29´n (bojesen-koefoed et al. 1999), the identification of large sedimentary basins and structures offshore, and high oil prices. a number of licensing rounds have been held with the result that at present (2009) 11 blocks covering more than 125 000 km2 have been awarded to company consortia (fig. 57; f.g. christiansen, personal communication 2009). in some areas outside the licensing areas an ‘open door’ policy has been introduced in order to encourage data acquisition while accepting higher risk. the result of the last ten years’ exploration activity is that there is now a modern regional data coverage of the region between c. 62° and 76°n, including more than 50 000 km non-exclusive seismic data and also airborne magnetic and gravity data covering very large areas. south and south-west greenland (c. 57–62°n) the shelf south of 62°n is relatively narrow with a steep, locally unstable, slope towards the ocean floor in the labrador sea. seismic data coverage is sparse, but new data acquisition in the recently awarded licence blocks (fig. 57), combined with investigations within the scope of the danish continental shelf programme, will provide a greatly improved data base in coming years. with the limited existing data base, assessment of the potential for oil and gas and possible play types is very speculative. however, there are indications of mesozoic rifting in the few available seismic lines, and a 141 ma old onshore coast-parallel dyke swarm bears witness of the initiation of rifting in the earliest cretaceous (watt 1969; l.m. larsen et al. 2009). southern west greenland (c. 62–68°n) this region is the site of extensive and deep basins with thick successions of mesozoic–cenozoic sediments. knowledge of the region is based not only on extensive seismic, magnetic and gravity surveys and seabed sampling but also on the data obtained from five wells drilled in the late 1970s: nukik-1, 65°32´n, 54°46´w nukik-2, 65°38´n, 54°46´w kangâmiut-1, 66°09´n, 56°11´w ikermiut-1, 66°56´n, 56°35´w hellefisk-1, 67°53´n, 56°44´w and a sixth well drilled in 2000: qulleq-1,63°49´n, 57°27´w no well penetrated the deepest sediments in the region; the oldest sediments encountered being those at the base 88 of the qulleq-1 well which are of santonian age (chri stiansen et al. 2001). in consequence interpretation of the age and lithologies of the deepest sediments is based largely on analogies with the labrador shelf where many more wells have been drilled. a prerequisite for petroleum prospectivity is the presence of a good source rock. although none of the wells sampled a good source rock, the live oil showings in vuggy paleocene basalts on the nuussuaq peninsula prove that such rocks exist in central west greenland (chri stiansen et al. 1996; bojesen-koefoed et al. 1999). the source rocks most likely to occur offshore southern west greenland are 1) lacustrine mudstones and coals in the lower cretaceous syn-rift kitsissut and appat sequences, 2) cenomanian–turonian organic-rich mudstones at or near the base of the post-rift kangeq sequence. these are correlatives of the marine cenoman ian–turonian source rock interpreted as having given rise to the itilli oil type, one of the five oil types occurring in the live oil showings on nuussuaq (bojesen-koefoed et al. 1999), 3) paleocene deltaic mudstones, correlatives of the creta ceous–paleocene source rocks in central west greenland from which three of the five oil types in live showings was derived (bojesen-koefoed et al. 1999) and 4) mudstones deposited distally relative to prograding palaeogene sequences that were described by dalhoff et al. (2003) the main play type involves block-faulted and tilted reservoir sandstones of the kitsissut and appat sequences with oil (and/or gas) derived from cenomanian–turonian source rocks and sealed by cenomanian–campanian mudstones of the kangeq formation or paleocene mudstones that drape the fault blocks (chalmers et al. 1993). anticlinal structures generated locally by transpression along the ungava fracture zone (fig. 48b) provide another potential for traps. new interpretations of the paleocene–mid-eocene seismic sequences combined with stratigraphic correlations with well data have shown that within the palaeogene succession there are sandy basinfloor fans and turbidite channel complexes encased in basin mudstones that could act as stratigraphic traps (dalhoff et al. 2003). from the kangâmiut-1 well wet gas (up to c5) was reported by chalmers et al. (1995), but the drill stem test produced only water from the drilling mud, not formation fluid (skaarup 2007). thus the possibility remains that a significant, untested, hydrocarbon field exists in the kangâmiut structure. central west greenland (c. 68–73°n) in this region palaeogene basalts are widespread which hampers interpretation of early paleocene and older sediments. however, since the 1:2 500 000 map was printed in 1995, a wealth of new data has been acquired in this region, allowing greatly improved interpretations and the compilation of an overview map showing the main structural elements in the area (fig. 58; gregersen et al. 2007). interest in acquiring new data in this area was stimulated by the discovery of live oil seeps in the adjacent onshore area (christiansen et al. 1996; bojesen-koefoed et al. 1999). five oil types have been identified in these seeps, three likely to have been derived from cretaceous and paleocene prodelta source rocks and a fourth, the itilli oil type, from dysoxic to anoxic marine shales, probably of cenomanian–turonian or older age. all these source rocks can be expected to occur in the offshore region. the new data have clearly revealed the existence of a number of deep basins with mesozoic and cenozoic sedimentary successions and also several large structures and 89 50 km 70°30'n 69°00'n 60°w 57°w 54°w nuussuaq basin nuussuaq disko gneiss ridge disko high disko lulissat high ilulissat graben aasiaat basin aasiaat structural trend k an ge rlu k s tr uc tu re hellefisk structure nagssugtôq subbasin sisimiut basin ik er m iu t f z ikermiut basin davis strait high n. ungava basin ilu lis sa t g ra be n e dg e gro#3 hellefisk-1 ikermiut-1 trend graben subbasin basin high ikermiut fault zone fault exploration well border with canada fig. 58. structural elements offshore central west greenland. from gregersen et al. 2007. qulleq-1 closures that could provide traps (gregersen & bidstrup 2008). the structures were initiated during early to midcretaceous rifting. syn-rift sediments deposited during this phase are likely to be sandstones with potential reservoir properties. during the subsequent late cretaceous quiet phase, a thick basinal mudstone unit was deposited, the equivalent of the kangeq sequence farther south. this may well contain source rocks that correlate with the source of the itilli oil type. renewed tectonic activity in the latest cretaceous – early paleocene caused uplift and formation of large structures, some of which could provide traps for hydrocarbons. prodelta paleocene source rocks equivalent to the suggested sources of three of the oil types identified in onshore seeps could be pres ent. late paleocene – eocene transpression related to the ungava fracture zone (fig. 49b) led to the formation of anticlinal structures that have a potential as traps. finally, during the eocene and especially during the late miocene and pliocene, the offshore basins subsided rapidly, and large sedimentary wedges prograded towards the west and south. direct hydrocarbon indicators (dhis) such as bright spots in some seismic lines are encouraging signs that this segment of greenland waters hosts live petroleum systems (gregersen et al. 2007). north-west greenland (73–77°n) the region is the site of some of the largest structures and deepest rift basins anywhere offshore west greenland (fig. 53; whittaker et al. 1997). this was first shown by the results of the kanumas reconnaissance seismic survey carried out in 1992; later public domain surveys carried out in north-east baffin bay in 2000 have provided a more detailed picture of this part of the rift system (gregersen 2008). in addition, recent gravity and magnetic surveys have supplemented knowledge of the deep basins and structural highs. no wells have been drilled in this region, so interpretation of the age and character of the sedimentary fill of the basins is based on analogies with onshore areas in north-east canada (bylot island) and central west greenland (the nuussuaq basin), and with the labrador sea. the bulk of the up to 12 km thick sedimentary fill is likely to be of cretaceous–neogene age. rifting probably started in the early cretaceous (whittaker et al. 1997), and sandstones deposited in the early syn-rift stage could be good reservoirs for hydrocarbons. during subsequent thermal subsidence a transgressive unit was deposited that is analogous to the latest cenoma n ian– turonian kanguk formation in the canadian arctic islands. near the base of this formation there are oilprone, marine shale source rocks which, however, in this area are thermally immature (núñez-betulu 1993). if the analogy to the kanguk formation holds, a similar marine source rock can be expected near the base of the transgressive unit in north-east baffin bay. support for this suggestion has been obtained by submitting samples of the source rock shales in the kanguk formation to hydrous pyrolysis. this yielded bitumen that shares a number of important characteristics with the itilli oil type occurring in vuggy basalts and fractures onshore central west greenland (bojesen-koefoed et al. 2004). the itilli oil type was generated from marine source rocks of presumably cenomanian–turonian age (boje sen-koefoed et al. 1999). marine source rocks usually have a wide areal distribution, so this source rock may well occur offshore north of 73°n and even tie up physically with the kanguk formation source rocks. the mudstones of the transgressive unit could also provide a seal to hydrocarbons trapped in the underlying sandstones in tilted fault blocks and anticlinal inversion structures. three of the other oil types described by bojesenkoefoed et al. (1999) were derived from cretaceous and paleocene prodelta source rocks. similar source rocks could also occur locally in north-east baffin bay. hydrocarbons generated from these source rocks could be trapped in the cores of anticlines formed during inversion and transpression, particularly in the northern part of the area. reservoir could be provided by depositional systems such as turbidite fan lobes shed off the inversion highs (whittaker et al. 1997). western north greenland (north of 80°n) lincoln sea (lincoln hav on the map) north of north greenland and ellesmere island (canada) contains an extensive shelf region (almost 500 × 200 km) with water depths below 500 m. the sea is normally covered by thick multi-year sea ice, and our present knowledge of the subsurface geology stems from two recent seismic refraction profiles (dahl-jensen et al. 2006) and earlier magnetic and gravity data. interpretation of the geophysical data suggests that the basin underlying the lincoln sea comprises a sedimentary sequence with a thickness of more than10 km. based on modelling from the seismic data it is assumed that the sedimentary succession in the lincoln sea basin is comparable with the deposits in the mesozoic–cenozoic sverdrup basin of arctic canada. by comparison with 90 known petroleum indications in the sverdrup basin it is inferred that strata in the lincoln sea basin may contain source rocks of marine shales of mid-triassic – latest jurassic and also marine type 2 source rocks of upper jurassic age (k. sørensen, personal communication 2009). mesozoic reservoir rocks may be widespread and intervening shales could form potential seals. acknowledgements the map was compiled in 1994–95 and printed in 1995. the onshore areas of the map sheet were compiled by j.c. escher with contributions from north and east greenland by h.-j. bengaard and n. henriksen. the offshore areas were compiled by t.c.r. pulvertaft, and the final compilation and legend design was carried out by j.c. escher (see also map legend). this second revised edition of the descriptive text was prepared in 2009. comments to the various sections have been provided by: the greenland precambrian shield: a.a. garde the gardar province: l.m. larsen thule supergroup: p.r. dawes carbonatites, kimberlites and associated rocks: l.m. larsen and t.f.d. nielsen the franklinian basin: j.r. ineson devonian to cenozoic basins in north-east greenland: s. piasecki cretaceous–palaeogene basins in west greenland: g. dam, g.k. pedersen and m. sønderholm palaeogene volcanic rocks and intrusions: l.m. larsen and t.f.d. nielsen quaternary sediments and glaciology: o. bennike offshore geology: t. dahl-jensen, t. funck and ø. engen mineral deposits: k. secher petroleum potential: f.g. christiansen, j.a. chalmers, u. gregersen, k. sørensen and n. skaarup. valuable comments and suggestions from other colleagues at the survey are also gratefully acknowledged. finally, the bulletin has benefited from thorough reviews by c.r.l. friend (uk), h. stendal (greenland) and g.n. oakey (canada), whose suggestions for improvements of text and illustrations have been incorporated with thanks. 91 92 allaart, j.h. 1975: geological map of greenland, 1:500 000, sheet 1, sydgrønland. copenhagen: geological survey of greenland. andersen, j.c.ø., rasmussen, h., nielsen, t.d.f. & rønsbo, j.g. 1998: the triple group and the platinova gold and palladium reefs in the skaergaard intrusion: stratigraphic and petrographic relations. economic geology 93, 488–509. andersen, s., bailey, j.c. & bohse, h. 1981: zr-y-u stratigraphy of the karkortokite–lujavrite sequence, southern ilímaussaq intrusion. in: bailey, j.c., larsen, l.m. & sørensen, h. 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[2] pearya – mainly exotic terrane. unit occurs only in northernmost ellesmere island, canada (p. 49). [3]–[5] middle mesoproterozoic to early neoproterozoic thule supergroup (p. 40), north-west greenland. age shown on map legend must be revised according to more recent acritarch studies. [3] carbonate and siliciclastic sediments: narssârssuk group, northwest greenland (p. 40). [4] shales and siltstones: dundas group, north-west greenland. [5] sandstones and shales: smith sound, nares strait and baffin bay groups, north-west greenland (p. 40). [6] palaeogene tholeiitic lavas, central west greenland (p. 60). [7] palaeogene picritic lavas, central west greenland (p. 60). [8] cretaceous–paleocene sediments: nuussuaq group, central west greenland (p. 57). [9] ordovician limestone in fault block in archaean gneiss. ‘fossilik’ locality (stouge & peel 1979), southern west greenland (65°25´n) (pp. 71, 75). [10] phanerozoic limestones in fault block within reworked archaean gneiss (peel & secher 1979), southern west greenland (66°32´n). [11] basaltic lavas: eriksfjord formation, mesoproterozoic, gardar province, south greenland (p. 38). [12] continental sandstones and conglomerates: eriksfjord formation, mesoproterozoic, gardar province, south greenland (p. 38). [13] pliocene–pleistocene sand-silt deposits: kap københavn formation, central north greenland (p. 62). [14] paleocene–eocene fluviatile and marine sandstones: thyra ø formation, wandel sea basin, eastern north greenland (pp. 54, 55). [15] upper cretaceous sandstones and shales: herlufsholm strand formation and correlatives, wandel sea basin, central and eastern north greenland (p. 54). [16] lower paleocene basic volcanics and volcanogenic sediments: kap washington group, wandel sea basin, central north greenland (p. 54). [17] upper jurassic – lower cretaceous sandstones and shales: ladegårdsåen formation and correlatives, wandel sea basin, central and eastern north greenland (p. 54). [18] upper permian – middle triassic shales and sandstones: trolle land group, wandel sea basin, central and eastern north greenland (p. 54). [19] upper carboniferous – lower permian carbonates: mallemuk mountain group, wandel sea basin, central and eastern north greenland (p. 54). [20] lower carboniferous sandstones and siltstones: sortebakker formation, wandel sea basin, eastern north greenland (p. 54). [21] silurian carbonates deposited on shelf and slope areas: washington land group, franklinian basin, north greenland (pp. 46, 47). [22] silurian sandstones and siltstones deposited in deep-water turbiditic trough: peary land group, franklinian basin, north greenland and ellesmere island (pp. 46, 47). [23] lower cambrian – lower silurian carbonates from shelf and slope areas: brønlund fjord, tavsen iskappe, ryder gletscher and morris bugt groups, franklinian basin, north and north-west greenland and ellesmere island (pp. 45, 46). [24] upper part of lower cambrian – lower silurian mudstones and shales: starved slope and trough deposits, vølvedal and amundsen land groups, franklinian basin, north greenland and ellesmere island (pp. 46, 82). [25] lower cambrian carbonates and siliciclastic sediments; shallowwater deposits: portfjeld and buen formations, franklinian basin, north and north-west greenland and ellesmere island (pp. 46). [26] upper neoproterozoic – lower cambrian calcareous mudstones and sandy turbidites: skagen, paradisfjeld and polkorridoren groups, deep-water trough deposits in the franklinian basin in north greenland and on ellesmere island (pp. 45, 46). [27] neoproterozoic siliciclastic and carbonate sediments: hagen fjord group, north greenland (p. 41). [28] upper neoproterozoic (marinoan ~ uppermost cryogenian) diamictites and sandstones: morænesø formation, central north greenland (pp. 41, 43). [29] neoproterozoic sandstones in caledonian nappe units: rivieradal group, in part older than the hagen fjord group, eastern north greenland (p. 41). [30] mesoproterozoic tholeiitic basalts (1380 ma): zig-zag dal basalt formation, central and eastern north greenland west of danmark fjord (pp. 33, 35). basalts in kronprins christian land, east of danmark fjord, indicated on the map as zig-zag dal basalt formation, are 1740 ma old, i.e. much older. [31] palaeoto mesoproterozoic sandstones: independence fjord group, central and eastern north greenland. shown as meso proterozoic on the map legend, but new age dating shows that parts are around 1740 ma (pp. 33, 35). [32] palaeogene (paleocene–eocene) plateau basalts: north-east greenland (p. 61). [33] upper jurassic and lower cretaceous shallow marine sandstones: raukelv, hesteelv, lindemans bugt and palnatokes bjerg formations and aptian–albian sediments, central east and north-east greenland (pp. 56, 57). a revised stratigraphy has been proposed by surlyk (2003). [34] middle–upper jurassic marine sandstones and shales: vardekløft, olympen, hareelv and bernberg formations, central east and north-east greenland (pp. 56, 57 ). a revised stratigraphy has legend explanation geological units onshore: indicated by numbers [1]–[86] in the map legend. offshore: indicated by letters [a]–[g] and six ornamentations in the map legend. 111 been proposed by surlyk (2003). [35] upper triassic – lower jurassic lacustrine sandstones and shales: kap stewart and neill klinter groups, central east greenland (p. 56). a revised stratigraphy has been proposed by surlyk (2003). [36] lower–upper triassic alluvial sandstones and lacustrine dolomites and shales: pingo dal, gipsdalen and fleming fjord formations, central east greenland (pp. 56, 57). [37] upper permian – lower triassic shallow marine carbonates, sandstones and shales: foldvik creek group and wordie creek formation, central east greenland (pp. 56, 57). [38] carboniferous – lower permian fluvial sandstones and shales, central east and north-east greenland (pp. 53, 55). [39] middle–upper devonian continental siliciclastic sedimentary rocks: vilddal, kap kolthoff, kap graah and celsius bjerg groups, north-east and central east greenland (pp. 52, 53). [40] cambro-ordovician dominantly limestones and dolomites in the east greenland caledonian fold belt: kløftelv, bastion, ella ø, hyolithus creek, dolomite point, antiklinalbugt, cape weber, narwhale sound and heim bjerge formations, northeast greenland (p. 50). [41] tillites of supposed marinoan (uppermost cryogenian) age in isolated occurrences, central east greenland (p. 43). [42] diamictites, sandstones, shales and dolostones in succession of marinoan – ediacaran age: tillite group in east greenland caledonian fold belt, north-east greenland (pp. 42, 43). [43] succession of siliciclastic, calcareous and dolomitic sediments of neoproterozoic (cryogenian) age: upper eleonore bay supergroup including lyell land, ymer ø and andrée land groups, east greenland caledonian fold belt, north-east greenland (pp. 42, 43). [44] succession of sandstones and siltstones of neoproterozoic (tonian–cryogenian) age: lower eleonore bay supergroup – the nathorst land group, east greenland caledonian fold belt, north-east greenland (pp. 42, 43). [45] palaeoproterozoic to early palaeozoic. mixed basement gneisses, metasedimentary rocks, greenstones, tillites and sedimentary rocks. exposed in tectonic windows in the east greenland caledonian fold belt, north-east greenland (pp. 31, 43). [46] late mesoproterozoic to early neoproterozoic metasedimentary rocks in the east greenland caledonian fold belt: krummedal supracrustal sequence and correlative smallefjord sequence, northeast greenland (pp. 38, 39). [47] eocene siliciclastic sedimentary rocks overlying lower palaeogene basalts, central and southern east greenland (p. 62). [48] eocene tholeiitic plateau basalts in central and southern east greenland (p. 61). [49] paleocene tholeiitic basalts with picritic intervals, southern east greenland (p. 61). [50] lower cretaceous – upper paleocene sandstones and shales. pre-basaltic succession in kangerlussuaq basin, southern east greenland (p. 60). [51] neogene and quaternary volcanic rocks in iceland, predominantly basalt lavas. [52] migmatites and gneisses of palaeoproterozoic to early neo proterozoic and caledonian origin in the east greenland caledonian fold belt, central east and north-east greenland (pp. 39, 50) and ellesmere island, canada (p. 49). [53] palaeogene felsic intrusions in east greenland (p. 62). [54] lateto post-kinematic granitic s.l. intrusions in the east greenland caledonian fold belt, central east and north-east greenland (pp. 39, 51). [55] neoproterozoic augen granite intrusions, deformed during the caledonian orogeny, central east and north-east greenland (pp. 38, 39). [56] mesoproterozoic intrusive complexes, mainly syenites: gardar province, south greenland (pp. 38, 82). [57] palaeogene mafic to intermediate intrusive complexes in east greenland (p. 62). [58] upper cretaceous gabbroic intrusion. pearya terrane, ellesmere island, canada. [59] middle jurassic carbonatite complex: qaqarssuk, southern west greenland (pp. 44, 82). [60] silurian pyroxenitic intrusions in archaean granulite gneisses: batbjerg complex, southern east greenland (p. 51). [61] late neoproterozoic carbonatite complex in archaean gneisses: sarfartoq, southern west greenland (pp. 44, 82). [62] palaeoproterozoic metasedimentary rocks (marbles and siliciclastic rocks) in the rinkian fold belt: karrat group comprising the mârmorilik, qeqertarssuaq and nûkavsak formations, central west greenland (pp. 25, 26, 81). [63] palaeoproterozoic basic metavolcanic rocks: sortis group in the northern border zone of the ketilidian orogen, south-west greenland (pp. 28, 29). [64] palaeoproterozoic metasedimentary rocks: vallen group in the northern border zone of the ketilidian orogen, south-west greenland (pp. 28, 29). [65] palaeoproterozoic acid metavolcanic rocks in the ketilidian orogen, south greenland (p. 30). [66] archaean acid metavolcanic rocks north-east of disko bugt in the rinkian fold belt, central west greenland (p. 26). [67] palaeoproterozoic, high-grade supracrustal units (paragneisses, marbles, quartzites and basic metavolcanic rocks) in palaeo proterozoic orogenic belts (pp. 23, 24, 28, 30, 32). [68] mesoand neoarchaean supracrustal rocks (amphibolites and gneissic metasediments) in the archaean craton, west greenland and south-east greenland (pp. 19, 23, 26). [69] eoarchaean supracrustal rocks (isua and akilia assemblages) in the archaean craton, southern west greenland (pp. 17, 18, 80). [70] palaeoproterozoic amphibolite facies gneisses (generally orthogneisses) dominantly of juvenile proterozoic origin. ketil idian orogen, south and south-east greenland (pp. 28, 29) and basement in northern part of caledonian fold belt in north-east greenland (pp. 31, 32, 38, 50). [71] palaeoproterozoic gneisses in granulite facies: inglefield land orogenic belt, north-west greenland and ellesmere island, canada (p. 28); nagssugtoqidian orogen, south-east greenland (p. 25) and caledonian fold belt in north-east greenland (p. 32). [72] mesoand neoarchaean orthogneisses in amphibolite facies. archaean craton, southern west greenland and south-east greenland (pp. 19, 20). [73] mesoand neoarchaean orthogneisses in granulite facies. archaean craton, southern west greenland and south-east greenland (pp. 19, 20). 112 [74] reworked amphibolite facies archaean gneisses in palaeo proterozoic orogens in west and south-east greenland (pp. 22, 25–27) and in the basement of the southern part of the east greenland caledonian fold belt, central east greenland (pp. 31, 32, 50). [75] reworked archaean granulite facies gneisses in palaeoprotero zoic orogens in central and southern west greenland (pp. 22, 24) and in south-east greenland (p. 24). [76] eoarchaean gneisses in the core of the archaean craton in southern west greenland: ‘amîtsoq gneiss’ (p. 18). [77] palaeoproterozoic rapakivi ‘granites’ in the ketilidian orogen, south greenland (pp. 28, 30, 38). [78] palaeoproterozoic granites: the julianehåb batholith in the ketil idian orogen, south greenland (pp. 28, 29); the prøven igneous complex in the rinkian fold belt (pp. 26, 27 ) and some granites in the nagssugtoqidian orogen of southern west and southeast greenland (pp. 24, 25) and within the basement of the caledonian fold belt in east greenland (p. 31). [79] neoarchaean post-tectonic granite complex: qôrqut granite, southern west greenland (p. 21). [80] mesoto neoarchaean granitic to tonalitic plutonic rocks; early–late kinematic intrusions: taserssuaq tonalite, ilivertalik augen granite, southern west greenland (pp. 20, 21, 24). in south-east greenland syenitic and granitic rocks (p. 21), and an intrusive complex in central west greenland (p. 26). [81] palaeoproterozoic intermediate plutonic rocks in the nagssugtoqidian orogen: arfersiorfik quartz diorite at 68°n, west greenland (p. 22, 24); ammassalik intrusive complex and similar rocks in se greenland (p. 24, 25). other occurrences in the ketilidian orogen in south greenland (p. 29) and in the caledonian fold belt in north-east greenland (p. 32). [82] mesoand neoarchaean post-tectonic intermediate and mafic intrusions in south-east and west greenland (pp. 21, 26) and north-west greenland (p. 27). [83] neoarchaean alkaline intrusive complex: skjoldungen alkaline province, south-east greenland (p. 21). [84] mesorchaean carbonatite sheet: tupertalik, southern west greenland (pp. 21, 44). [85] mesoarchaean anorthositic rocks in the archaean craton: fiskenæsset complex and correlatives, southern west greenland (pp. 19, 80); also in central west greenland (p. 26), in southeast greenland (p. 24) and in the thule region (c. 77°30´n) north-west greenland (p. 27). [86] palaeoproterozoic gabbro-anorthosite, east greenland caledo nian fold belt, north-east greenland (76°n) (stecher & henriksen 1994). offshore [a] areas underlain by continental crust with or without cover of sedimentary rocks and tertiary volcanic rocks (p. 66). [b] transition zone between continental and oceanic crust. in many areas thought to consist of continental crust with increasing intensity of dykes and intrusions as oceanic crust is approached (p. 66). off south-west greenland transition zone is extremely thin continental crust flanked to the south-west by a zone of serpentinised mantle peridotite. [c]–[f ] areas underlain by oceanic crust, divided according to age at 15 million year intervals. oldest oceanic crust [f] was formed more than 45 million years ago. divisions based on sea-floor spreading magnetic anomalies (p. 66). [g] oceanic crust of unspecified age (pp. 66, 70). ornamentations palaeogene volcanic rocks at seabed or concealed, latter only shown in areas underlain by continental crust: north-east greenland 72–75°n (p. 74); west and north-west greenland 68–73°n (p. 76). buried volcano with high relief, central east greenland, 69°n. intrusions in sedimentary and volcanic rocks, east greenland (71–73°n). probably of palaeogene age (p. 74). areas with widespread salt deposits of supposed late palaeozoic age, north-east greenland shelf 76°30´–79°30´n (p. 73). sedimentary basins with thicknesses over 4 km (pp. 73–77). most sediments are of late palaeozoic – cenozoic age. little known basins with thick sedimentary successions (pp. 73, 74, 76, 77). 113 a aasiaat 68°43′/52°53′ 6 akia 64°24′/51°43′ 7 alert (canada) 82°30′/62°09′ 5 alluitsup paa 60°28′/45°34′ 7 ameralik [lysefjord] 64°07′/51°00′ 7 ammassalik 65°36′/37°38′ 10 anap nunaa 69°57′/50°30′ 6 andreé land 73°42′/26°25′ 12 ardencaple fjord 75°20′/21°00′ 12 arfersiorfik 68°10′/52°28′ 6 arsuk 61°11′/48°26′ 7 atammik 64°48′/52°12′ 7 attu 67°57′/53°38′ 6 b bache peninsula (canada) 79°13′/76°50′ 5 baffin bugt 73°00′/62°00′ 6 bessel fjord 75°59′/21°00′ 11 bildsøe nunatakker 78°08′/23°48′ 11 bjørnesund 62°55′/50°10′ 7 blosseville kyst 68°49′/26°00′ 12 bredefjord 60°55′/46°25′ 7 breitafjör›ur (iceland) 64°25′/23°00′ 13 c camp century 77°11′/61°07′ 5 canning land 71°40′/22°15′ 12 carey øer 76°43′/72°58′ 5 charcot land 71°53′/29°45′ 12 christianshåb 68°49′/51°11′ 6 clavering ø 74°18′/21°00′ 12 constable pynt 70°45′/22°36′ 12 d daneborg 74°19′/20°14′ 12 danell fjord [iluileq] 60°53′/43°08′ 10 danmark fjord 81°10′/21°30′ 11 danmark stræde 66°10′/27°00′ 12, 13 danmarkshavn 76°46′/18°39′ 11 davis stræde 68°00′/57°00′ 6 disko 69°45′/53°30′ 6 disko bugt 69°11′/52°45′ 6 dome grip (summit) 72°35′/37°38′ 9 dove bugt 76°37′/20°00′ 11 dronning louise land 76°30′/24°30′ 11 dye 3 65°11′/43°50′ 10 e egedesminde 68°43′/52°53′ 6 eleonore bugt 73°26′/25°23′ 12 ella ø 72°55′/25°05′ 12 ellesmere island (canada) 80°00′/80°00′ 5 f fiskefjord 64°54′/51°33′ 7 fiskenæsset 63°05′/50°41′ 7 frederick e. hyde fjord 83°10′/30°30′ 8 frederikshåb 62°00′/49°40′ 7 frederikshåb isblink 62°35′/49°55′ 7 freuchen land 82°20′/43°30′ 8 g gåsefjord 70°04′/28°00′ 12 geikie plateau 69°56′/25°30′ 12 germania land 77°06′/18°55′ 11 giesecke isfjord 73°36′/55°58′ 6 gisp 2 72°35′/38°27′ 9 gletscherland 72°40′/27°00′ 12 godhavn 69°15′/53°33′ 6 godthåb 64°11′/51°45′ 7 godthåbsfjord 64°25′/51°25′ 7 grænseland 61°23′/47°53′ 7 greely fiord (canada) 80°24′/83°00′ 5 grønlandshavet 77°00′/10°00′ 11 grønnedal 61°14′/48°06′ 7 gunnbjørn fjeld 68°51′/29°52′ 12 h hagen fjord 81°35′/25°30′ 8 hall bassin 81°30′/63°00′ 5 hall bredning 70°54′/24°45′ 12 hall land 81°30′/60°00′ 5 hans ø 80°50′/66°38′ 5 hans tausen iskappe 82°32′/38°00′ 8 hareøen 70°26′/54°55′ 6 hellefisk-1 67°52′/56°44′ 6 herluf trolle land 82°30′/26°30′ 8 place names register includes all place names shown on the geological map. the names in square brackets are some well-known alternative names that do not appear on the map. map segment numbers refer to the index map on page 10 (fig. 1). in the alphabetical sorting the danish letters æ, ø and å are treated as ae, o and a; for convenience øfjord also follows z. place name lat. / long. map n w segment place name lat. / long. map n w segment 114 hochstetter forland 75°30′/19°53′ 12 hold with hope 73°44′/21°10′ 12 holm land 80°26′/17°30′ 11 [holm ø] see kiatassuaq holsteinsborg 66°56′/53°40′ 6 hovgaard ø 79°55′/18°30′ 11 hudson land 73°49′/23°00′ 12 humboldt gletscher 79°30′/63°30′ 5 i ikaasakajik 70°55′/27°00′ 12 ikeq 64°56′/40°35′ 10 [ikerasassuaq] see prins christian sund ikermiut-1 66°56′/56°35′ 6 ikersuaq 60°55′/46°25′ 7 ikertivaq 65°29′/39°35′ 10 île de france 77°49′/17°50′ 11 ilimananngip nunaa 70°43′/26°48′ 12 illoqqortoormiut 70°29′/21°58′ 12 illorsuit 71°09′/53°40′ 6 [iluileq] see danell fjord ilulissat 69°13′/51°07′ 6 independence fjord 82°05′/29°30′ 8 inglefield land 78°44′/69°00′ 5 ingolf fjord 80°30′/18°00′ 11 innaanganeq 75°55′/66°28′ 5 inuit qeqertaat 83°40′/30°35′ 8 ísafjör›ur (iceland) 66°05′/23°10′ 13 island (iceland) 65°00′/18°00′ 13 isukasia 65°11′/49°48′ 7 ittertivaa 69°25′/24°06′ 12 ivisaartoq 64°49′/49°58′ 7 ivittuut 61°12′/48°10′ 7 j jakobshavn 69°13′/51°07′ 6 jakobshavn isfjord 69°10′/50°30′ 6 jameson land 71°10′/23°15′ 12 j.c. christensen land 81°40′/29°30′ 8 johannes v. jensen land 83°20′/32°00′ 8 jøkelbugten 78°38′/20°00′ 11 j.p. koch fjord 82°45′/44°30′ 8 julianehåb 60°43′/46°03′ 7 k kaffeklubben ø 83°40′/30°35′ 8 kane bassin 79°30′/69°00′ 5 kangaamiut 65°50′/53°21′ 7 kangaarsugsuaq 77°01′/71°23′ 5 kangaatsiaq 68°19′/53°28′ 6 kangâmiut-1 66°09′/56°11′ 7 kangeq 61°50′/42°06′ 10 kangerlussuaq (east greenland) 68°22′/32°12′ 9, 12 kangerlussuaq (west greenland) 66°24′/52°30′ 7 [kangerlussuatsiaq] see lindenow fjord kangersik kiatteq 71°30′/26°00′ 12 kangersuatsiaq 72°23′/55°34′ 6 kangertittivaq 70°17′/23°00′ 12 kangertittivatsiaq 66°21′/35°43′ 10 kangikajik 70°09′/22°03′ 12 kap alexander 78°11′/73°02′ 5 kap brewster 70°09′/22°03′ 12 kap bryant 82°20′/55°15′ 8 kap cort adelaer 61°50′/42°06′ 10 kap dalton 69°25′/24°06′ 12 kap edvard holm 67°51′/32°11′ 9 kap eiler rasmussen 82°35′/19°45′ 8 kap farvel 59°47′/43°55′ 10 kap franklin 73°15′/22°10′ 12 kap gustav holm 66°34′/34°20′ 10 kap københavn 82°23′/20°57′ 8 kap morris jesup 83°39′/33°25′ 8 kap møsting 63°41′/40°31′ 10 kap parry 77°01′/71°23′ 5 kap ravn 68°26′/28°16′ 12 kap tordenskjold 61°24′/42°23′ 10 kap washington 83°33′/38°40′ 8 kap york 75°55′/66°28′ 5 karrat isfjord 71°34′/52°25′ 6 keflavik (iceland) 64°00′/22°30′ 13 kejser franz joseph fjord 73°21′/23°30′ 12 kennedy kanal 80°40′/68°00′ 5 kialiip imaa 66°55′/33°45′ 10 kiatassuaq [holm ø] 74°30′/57°00′ 6 kilen 81°11′/13°25′ 11 kitsissut 76°43′/72°58′ 5 kobberminebugt 60°55′/48°17′ 7 køge bugt 64°56′/40°35′ 10 kong oscar fjord 72°22′/24°00′ 12 kronprins christian land 80°40′/21°00′ 11 kuhn ø 74°50′/20°20′ 12 kullorsuaq 74°34′/57°10′ 6 kulusuk 65°34′/37°11′ 10 kuummiut 65°52′/37°01′ 10 l lake hazen (canada) 81°47′/70°50′ 5 lambert land 79°19′/20°48′ 11 lauge koch kyst 76°20′/60°00′ 5 lincoln hav 83°25′/57°00′ 8 lindenow fjord [kangerlussuatsiaq] 60°30′/43°30′ 10 liverpool land 70°55′/22°00′ 12 lyell land 72°38′/25°35′ 12 [lysefjord] see ameralik m maarmorilik 71°08′/51°18′ 6 mallemukfjeld 80°10′/17°04′ 11 place name lat. / long. map n w segment place name lat. / long. map n w segment 115 maniitsoq 65°25′/52°52′ 7 melville bugt 75°45′/60°50′ 5 mestersvig 72°14′/23°55′ 12 midternæs 61°37′/47°56′ 7 milne land 70°43′/26°48′ 12 mont forel 66°56′/36°49′ 10 mylius-erichsen land 81°00′/26°00′ 8 n nakkehoved 81°42′/13°03′ 11 nanortalik 60°09′/45°15′ 7 nansen fjord 68°17′/29°50′ 12 nansen land 82°56′/44°20′ 8 nansen sound (canada) 81°00′/90°00′ 5 nares stræde 80°00′/69°00′ 5 narsaq 60°55′/46°03′ 7 narsarsuaq 61°10′/45°25′ 7 nassuttooq 67°45′/53°00′ 6 neriap nunaa 61°23′/47°53′ 7 nertiit kangersivat 70°04′/28°00′ 12 niaqornaarsuk 68°14′/52°52′ 6 niaqorsuaq 75°04′/58°05′ 6 nioghalvfjerdsfjorden 79°34′/21°00′ 11 nordatlanten 62°00′/29°00′ 13 nordlandet 64°24′/51°43′ 7 nordostrundingen 81°21′/11°20′ 11 nordre strømfjord 67°45′/53°00′ 6 nordvestfjord 71°30′/26°00′ 12 norske øer 79°07′/17°50′ 11 nukik-1 65°31′/54°45′ 7 nukik-2 65°38′/54°46′ 7 nunakuluut [nunarssuit] 60°46′/47°57′ 7 nunap isua 59°47′/43°55′ 10 [nunarssuit] see nunakuluut nuuk 64°11′/51°45′ 7 nuup kangerlua 64°25/51°25′ 7 nuussuaq 70°35′/52°55′ 6 nyeboe land 81°45′/57°00′ 8 o øfjord 70°55′/27°00′ 12 p paamiut 62°00′/49°40′ 7 peary land 82°35′/31°00′ 8 petermann bjerg 73°05′/28°37′ 12 petermann gletscher 80°35′/59°35′ 5 pituffik 76°33′/68°15′ 5 prins christian sund [ikerasassuaq] 60°07′/43°30′ 10 prinsen af wales bjerge 68°56′/32°30′ 9 prøven 72°23′/55°34′ 6 q qaanaaq 77°28′/69°14′ 5 qaqortoq 60°43′/46°03′ 7 qasigiannguit 68°49′/51°11′ 6 qeqertarsuaq (disko) 69°45′/53°30′ 6 qeqertarsuaq (godhavn) 69°15′/53°33′ 6 qeqertarsuatsiaat 63°05′/50°41′ 7 qeqertarsuatsiaq 70°26′/54°55′ 6 qeqertarsuup tunua 69°11′/52°45′ 6 qimusseriarsuaq 75°45′/60°50′ 5 qullissat 70°05′/53°01′ 6 qunaranaaq 61°24′/42°23′ 10 r ravn storø 62°43′/50°23′ 7 red head 75°04′/58°05′ 6 renland 71°20′/26°45′ 12 reykjavik (iceland) 64°10′/22°00′ 13 rink isbræ 71°47′/51°23′ 6 robeson kanal 81°53′/62°00′ 5 s saqqisikuik 63°22′/41°35′ 10 sarfartoq 66°25′/51°23′ 7 scoresby land 71°45′/25°00′ 12 scoresby sund 70°17′/23°00′ 12 scoresbysund 70°29′/21°58′ 12 sermersuaq (humboldt gletscher) 79°30′/63°30′ 5 sermersuaq (steenstrup gletscher) 75°17′/57°53′ 6 sermiligaarsuk 61°30′/48°40′ 7 sermilik 66°11′/37°36′ 10 shannon 75°08′/18°15′ 12 sherard osborn fjord 82°05′/52°05′ 8 sisimiut 66°56′/53°40′ 6 skærfjorden 77°30′/19°30′ 11 skærgårdshalvø 68°09′/31°45′ 12 skjoldungen 63°22′/41°35′ 10 skrækkens bugt 66°55′/33°45′ 10 smith sund 78°30′/74°00′ 5 snæfellsnes (iceland) 64°50′/23°37′ 13 søndre strømfjord (airport) 66°58′/50°57′ 6 søndre strømfjord (fjord) 66°24′/52°30′ 7 station nord 81°35′/16°41′ 11 stauning alper 72°00′/25°00′ 12 steenstrup gletscher 75°17′/57°53′ 6 store koldewey 76°30′/19°00′ 11 storstrømmen 76°53′/22°50′ 11 suess land 72°58′/25°35′ 12 sukkertoppen 65°25′/52°52′ 7 sullorsuaq 70°16′/53°25′ 6 svartenhuk halvø 71°45′/54°50′ 6 sydprøven 60°28′/45°34′ 7 t taartoq 61°25′/48°50′ 7 tasiilap karra 66°34′/34°20′ 10 place name lat. / long. map n w segment place name lat. / long. map n w segment 116 tasiilaq 65°36′/37°38′ 10 thule 77°28′/69°14′ 5 thule air base 76°33′/68°15′ 5 traill ø 72°40′/23°43′ 12 tuttut nunat 71°20′/26°45′ 12 u ubekendt ejland 71°09′/53°40′ 6 ullersuaq 78°11′/73°02′ 5 umiivik 64°16′/40°35′ 10 united states range (canada) 82°00′/72°00′ 5 upernavik 72°47′/56°10′ 6 upernavik isfjord 72°55′/55°30′ 6 ussing isfjord 73°54′/56°00′ 6 uummannaq 70°41′/52°08′ 6 v vaigat 70°16′/53°25′ 6 victoria fjord 82°09′/47°45′ 8 w waltershausen gletscher 74°09′/25°30′ 12 wandel dal 82°14′/33°30′ 8 wandel hav 82°30′/12°00′ 11 ward hunt ice shelf (canada) 83°08′/75°00′ 5 warming land 81°34′/52°50′ 8 washington land 80°30′/64°00′ 5 watkins bjerge 68°51′/29°30′ 12 wollaston forland 74°26′/19°35′ 12 wulff land 81°51′/48°30′ 8 y ymer ø 73°11′/24°30′ 12 ø øfjord 70°55′/27°00′ 12 place name lat. / long. map n w segment place name lat. / long. map n w segment 117 abbreviations fm formation gp group sgp supergroup ceg central east greenland cng central north greenland cwg central west greenland eg east greenland (includes ceg, neg, seg, seg) eng eastern north greenland neg north-east greenland ng north greenland nwg north-west greenland seg southern east greenland seg south-east greenland sg south greenland swg southern west greenland swg south-west greenland wg west greenland (includes cwg, nwg, swg, swg) wng western north greenland a aasiaat domain 23 aasiat basin: cwg 15 aasiat structural trend: cwg 89 aasivik terran: swg 19 accretion of precambrian crustal blocks: swg 15, 21, 48 aeromagnetic data 24, 66, 67 agatdal formation (paleocene): cwg 57, 58, 59 agpaitic rocks, see alkaline rocks akia terrane (mesoarchaean): swg 19 akilia association (eoarchaean): swg 18 albian: neg, cwg, seg, swg 56, 58, 60, 75 albian–campanian (cretaceous): cwg 57 alkali acid rocks (gardar): sg 35 alkaline rocks: eg, sg 62, 82 ameralik dykes (eoarchaean): swg 18, 18, 36 amikoq: swg 79 amitsoq: sg 79 amîtsoq gneiss (eoarchaean): swg 18, 36, 37 amîtsoq iron-rich suite (eoarchaean): swg 18 ammassalik intrusive complex (palaeoproterozoic): seg 24 ammassalik mobile belt (palaeoproterozoic): seg, seg 24 amundsen land gp (ordovician): ng 46, 46, 48, 78 anap nunâ gp (palaeoproterozoic): cwg 26 andrée land gp (neoproterozoic): ceg, neg 42, 43 anorthositic rocks cwg (mesoarchaean) 26 neg (proterozoic), see legend explanation [86] 112 nwg (neoarchaean) 27 seg (nagssugtoqidian orogen) 24 swg (mesoarchaean) 16, 19, 80 antiklinalbugt fm (ordovician): neg 50 antimony (sb): neg 79 appat sequence (cretaceous): swg 75, 76, 76, 89 appinites (proterozoic): sg, seg 29, 37 aptian–albian (cretaceous) neg 56 seg 60 archaean eoarchaean craton: swg 16, 17, 18, 36 craton: seg, swg, swg 14, 14, 15, 16, 17, 19–22, 21, 28, 36, 44, 79, 80 gneisses (meso–neoarchaean): swg, seg 19, 21 intrusive rocks (meso–neoarchaean): seg, swg 20, 21 mineral deposits 80, 81 reworked rocks 14, 14 ceg, neg 31 seg, seg, wg 21–28, 23, 25 supracrustal rocks eoarchaean: swg 17 meso–neoarchaean cwg 23 seg, swg, swg 19, 24 terrane boundaries 16, 17, 19–21 archaean basement beneath the inland ice 32 archaean orogenesis 15–20 archaean terraines 14 arco 86 arfersiorfik quartz diorite: swg 23, 24 arveprinsens ejland: cwg 79 asphalt residues 84 atâ igneous complex 26 ataata kuua: cwg 59 atammik: swg 22 atane fm (cretaceous): cwg 57, 58 atanikerdluk fm (paleocene): cwg 58 atlantic richfield company 86 attu: swg 79 augite syenite (gardar): sg 35 b bache peninsula (canada) 69, 72 baffin bay gp (meso–neoproterozoic): nwg 40, 40 baffin bay (bugt): nwg 71, 85, 90 baffin island (canada) 71 index in the listing the danish letters æ, ø and å are treated as ae, o and a. for the use of the geographical subdivisions see the map on page 4. page numbers in bold refer to figures. 118 baltica 50 banded iron formation isua supracrustal sequences: swg 17, 17, 80 cwg 26 nwg 27, 81 barden bugt fm (meso–neoproterozoic): nwg 40 barite: ng, ceg 79 barremian–albian (cretaceous), offshore: swg 75 basalts archaean 19 paleocene: cng 53, 54 palaeoproterozoic: swg 28 meso–neoproterozoic: ng, sg, nwg 33, 35, 37, 39 palaeogene 79, 85 cwg 37, 58, 59, 60, 89, 90 eg 37, 60, 61, 61, 62 tertiary offshore 67–69, 74, 76, 76, 85 base metals, exploration 9, 78 basement provinces 9, 14, 14 basic dykes, see dykes and sills basins, see sedimentary basins bastion fm (cambrian): neg 50 batbjerg complex (caledonian): seg 51 bathymetry 13 bernbjerg fm (jurassic): neg 56 beryllium, prospect / deposit: sg 79, 82 bgr (bundesanstalt für geowiss. und rohstoffe) 71 bitumen 90 bjørnesund block: swg 16 black angel (sorte engel) lead-zinc deposit: cwg 26, 79, 81 blosseville gp (cretaceous–palaeogene): seg 60 blosseville kyst basin: ceg, seg 75 petroleum potential 88 blyklippen lead-zinc deposit (mestersvig): ceg 79, 83 borgtinderne intrusion (tertiary): seg 62 bredehorn: ceg 79 brogetdal: neg 79 brønlund fjord gp (cambrian): ng 41, 46 buen fm (cambrian): ng 41, 45, 46, 49 bundesanstalt für geowiss. und rohstoffe (bgr) 71 c caledonides, east greenland 50, 51 archaean – palaeoproterozoic basement: ceg, neg 31, 31 sedimentary rocks of unknown age: neg 43 fold belt units: ceg, eng, neg 51 reworked neoproterozoic units: ceg, neg 38, 38, 39, 39 sedimentary successions (neoproterozoic–palaeozoic) 41–44, 43, 44, 50 intrusions: ceg, neg 51, 52 orogeny: ceg, eng, neg 50, 51, 52 cambrian sediments: ng 28, 41, 45, 46, 48, 49 deep-water trough: ng 45, 46, 48 source rocks: ng 84 cambrian–ordovician sediments: neg 43, 50, 50 cambrian–silurian sediments: ng, nwg 45, 46 campanuladal fm (neoproterozoic): cng, eng 42 cape camperdown fm (mesoproterozoic): nwg 40 cape combermere fm (mesoproterozoic): nwg 40 cape dyer (canada) 69 cape smith orogen (canada) 15 cape weber fm (ordovician): neg 50 carbonatite complexes 44 mesoarchaean: swg 21 gardar (mesoproterozoic): sg 38 neoproterozoic / jurassic: swg 44 carboniferous sediments ceg, neg 53, 53, 55, 55, 86, 87 eng 53, 54, 54, 79, 84, 86 ng 45, 48, 48 carboniferous–permian salt formation, offshore: neg 73 celestite: ceg 79 celsius bjerg gp (devonian–carboniferous): neg 53, 53 charcot land supracrustal succession (palaeoproterozoic): ceg 31 charcot land tectonic window: ceg 31 charnockitic intrusive rocks nwg (prøven ignous complex) 27 seg (ammassalik intrusive complex) 24 chromite mineralisation: swg 80 chromium: swg 79, 80 chron numbers (magnetochrons) 13, 66, 70 citronen fjord: cng 76, 78, 79, 80 citronen fjord lead-zinc deposit: cng 80 clarence head fm (mesoproterozoic): nwg 40 clavering ø: neg 79 climate, past 63, 65 coal: cwg 79 coal deposit (cretaceous): cwg 57, 58, 60 cobalt: cwg 79 continental break up, palaeogene 60, 62 continental crust, offshore 66–73, 68, 74, 75, 75 continental margin eg, ng 66–70, 73 wg 70–72 continental shelf 11, 63, 71, 86 continent–ocean boundary (cob) 66, 67 continent-ocean transition zone 67 copper: nwg, cwg, sg, seg, neg 79 cretaceous, sediments 79 ceg 55, 56, 56 cng, eng 48, 53, 54, 55 cwg 8, 57, 58, 83, 86 neg 55, 56, 57, 74, 74 offshore swg, swg 71, 75, 76, 86, 88–90 seg 60 ng, nwg 70, 77 cryolite: swg 78, 79, 80, 82 d danell fjord: seg 38 danish expeditions to east greenland 11 danmark stræde (denmark strait): seg 67 119 danmarks geologiske undersøgelse (dgu) 11 danmarkshavn basin: neg 74, 84, 85, 87 danmarkshavn ridge: neg 73, 74 davis stræde (davis strait): swg 69, 71, 75, 85 davis strait high: swg 69, 71, 75, 89 de geer megashear 70 deep drillings inland ice 65 onshore: cwg 57, 85, 89 offshore seg 66, 67, 75 swg 66, 68, 71, 84, 87 through inland ice 32 denmark strait fracture zone 67 devon dal: ceg 79 devonian sediments neg 51, 52, 53, 53, 79 ceg 55 wng 45, 48, 49 offshore: neg 73, 74, 87 diamictites (neoproterozoic) morænesø fm: cng 41, 41 tillite gp: neg 41, 42, 43, 43, 51 diamond: swg 79 diamond exploration: swg 44, 78 ‘diamond province’: swg 44 disko bugt suture: cwg 15, 22 disko bugt: cwg 4, 14, 15, 85 disko gneiss ridge: cwg 89 disko high: cwg 89 dolerite dykes, see dykes and sills dolomite point fm (cambrian): neg 50 dove bugt: neg 32 drill holes inland ice 32, 65 offshore: swg 66, 68, 71, 84, 87 onshore: cwg 57, 85, 89 dundas gp (meso–neoproterozoic): nwg 40, 40 dykes and sills ameralik (palaeoarchaean) 18, 18, 36, 37 archaean 18, 18, 36, 37 gardar (mesoproterozoic): sg 35, 37, 38 kangâmiut (palaeoproterozoic): swg 22, 36, 37 ‘md’ (palaeoproterozoic): swg, swg 37 midsommersø (mesoproterozoic): cng, eng 33, 34, 37, 40 phanerozoic 37 proterozoic 37 tarssartôq (palaeoarchaean): swg 37 ‘td’ (cretaceous): sg, swg 37 tertiary 37, 61, 62, 67 umîvik (palaeoproterozoic): seg, seg 37 e east greenland escarpment: ceg 75 eclogite: neg 32, 50 swg 24 eemian (quaternary) 63, 65, 65 eleonore bay sgp (neoproterozoic): ceg, neg 39, 41, 42, 43, 44, 51, 52 eleonore sø series (palaeoproterozoic): neg 43 eleonore sø window: neg 31 ella ø fm (cambrian): neg 50 ellesmere island (canada) 4, 33, 39, 45–49, 48, 72, 73, 85, 90 ellesmere-devonian terrane (canada) 15 ellesmerian fold belt (palaeozoic): ng 48, 49, 78, 85 ellesmerian orogeny (palaeozoic): ng 51 eoarchaean 15 supracrustal rocks 17 gneisses 16, 18, 18 eocene oceanic crust 68 eocene sediments offshore ceg 75, 75, 88 cwg 76, 90 swg 76, 76, 89 onshore: seg 62 eng 55 eocene triple junction 70 eqalulik formation (paleocene): cwg 58, 59 eqalussuit: swg 79 eqi: cwg 79 eriksfjord fm (gardar): sg 35, 38 etah gp (palaeoproterozoic): nwg 28 eurasia basin 70 eurekan orogen 72, 73 evaporites, onshore carboniferous–permian: neg 73 permian: ceg 56 ordovician: ng 45, 46 neoproterozoic: neg 42 meso–neoproterozoic: nwg 40 evaporites, offshore: neg 73 exotic terrane, ellesmere island (canada), see pearya 48, 49 exploration wells (hydrocarbon): wg 57, 66, 67, 71, 84, 87, 88, 88 f færingehavn terrane: swg 16, 17 færingehavn: swg 16 fiskefjord block: swg 16 fiskefjord: swg 20, 36 fiskenæsset: swg 79 fiskenæsset complex (mesoarchaean): swg 19, 20, 80 flammefjeld: seg 79 fleming fjord fm (triassic): ceg 56 foldvik creek gp (permian): ceg 56 ‘fossilik’ (ordovician): swg, see legend explanation [9] 110 foxe fold belt (canada) 27 foyaite (gardar): sg 35 fracture zones, oceanic: eg, wg 66, 67, 69, 69, 70–72, 74, 76, 89, 90 fram strait (greenland–svalbard) 69, 70 franklinian basin (cambro-silurian) ng, nwg 28, 41, 45, 46, 48, 49, 51 120 mineral deposits: ng 78, 79 petroleum potential: ng 84 franz joseph thrust sheet unit: neg 51 fylla fault complex: cwg 76 fyns sø fm (neoproterozoic): cng, eng 41, 42 g gakkel ridge: eng 69, 70, 73 garbenschiefer: swg 17 gardar province (mesoproterozoic): sg 15, 16, 35 intrusive complexes 35, 38, 44 sediments 38 sills and dykes 38 volcanics 38 mineral deposits 78, 79, 82 gas 84, 86–89 gåsefjord: ceg 61 gåseland window: ceg 43 geochemical surveys 11, 20, 21, 78 geodetic institute (gi) 13 geological survey of denmark (dgu) 11 geological survey of denmark and greenland (geus) 11 geological survey of greenland (ggu) 11 geophysical surveys 11, 66 gipsdalen fm (triassic): ceg 56 gisp 2 (drilling through the inland ice): neg 14, 32 glaciation neoproterozoic morænesø fm: cng 41 tillite gp: neg 42 quaternary 63 glaciers 12, 63, 64 glaciology 64 glaciomarine sediments (offshore) 67 godthåb (nuuk): swg 14, 16, 18 godthåbsfjord (nuup kangerlua): swg 17, 18, 18–21, 36, 80 godthåbsfjord-ameralik belt: swg 16 gold 79 gold deposits nalunaq: sg 78, 79, 81, 81 nuup kangerlua (godthåbsfjord): cwg 80 skaergaard: seg 62, 79, 83 gold exploration 78 gold occurrences (disko bugt): cwg 80 grænseland: sw 28, 29 granulite facies neg 32 nwg 27, 28 seg 25 seg 51 sg 30 wg 16, 17, 18, 20, 21, 22, 24 graphite: nwg, swg, sg 18, 23, 60, 79 gravimetric data 12, 66 greenland geological map 10, 85 size 13 ice-free land areas 13 greenland fracture zone: neg 74 greenland mineralisation data bank 80 greenland ridge: neg 69, 69, 70 greenland sea (grønlandshavet): neg 85 greenstone belts (archaean) disko bugt: cwg 26 eoarchaean: swg 18 grenvillian: ceg, neg 43, 51 gro#3 exploration well: cwg 58, 59, 86, 89 grønarctic energy inc. 86 grønlandshavet (greenland sea): neg 85 grønnedal: swg 79 h hagar bjerg thrust sheet: neg 38, 39, 51 hagen fjord gp (neoproterozoic): cng, eng 33, 35, 40, 41, 41, 42, 51 hagen fjord: eng 42 hall land: wng 46 hammer dal: cwg 79 harder bjerg fm: neg 53 harder fjord fault zone: cng 53 hareelv fm (jurassic): ceg 56, 87 hauge bjerge fm (silurian): ng 47 hecla high: swg 76, 76 heimbjerge fm (ordovician): neg 50 hellefisk structure: swg 89 hellefisk-1 exploration well: swg 76, 88 herlufsholm strand fm (cretaceous): cng, eng 54 hesteelv fm (cretaceous): ceg 56 hold with hope: neg 61, 62 holocene (quaternary) 63 hovgård ridge: neg 70 hyaloclastite (tertiary) cwg 59, 60 seg 61 hydrocarbon exploration, concession areas 87 hydrocarbon potential 9, 84, 87, 88, 90 hydrocarbons 87–90 hyolithus creek fm (cambrian): neg 50 i iapetus ocean 50 ice-core records 65 ika: swg 79 ikermiut basin: swg 85, 89 ikermiut fault zone: swg 89 ikermiut-1 exploration well: swg 71, 76, 88, 89 ikkattoq gneisses (mesoarchaean): cwg 20 ilimaussaq: sg 35, 79, 80, 82 ilímaussaq intrusion (gardar): sg 35 ilivertalik augen granite (mesoarchaean): swg 16, 21 illinoian (quaternary) 63 illorsuit: sg 79 ilukunnguaq: cwg 79 ilulissat graben: cwg 89 121 ilulissat high: cwg 89 imilik intrusion (tertiary): seg 62 independence fjord gp (palaeo–mesoproterozoic): cng, eng 33, 34, 35, 40, 51 independence fjord: cng 34 inglefield orogenic belt (palaeoproterozoic): nwg 14, 14, 15, 21, 27 ingolf fjord: eng 34 inland ice bedrock below 32 former extent 63, 63, 65 gisp core 14, 32 mass balance 64, 65 size 13 thickness 64 volume 64 iron: nwg, cwg, swg, swg, neg 79 iron deposit banded iron formation (archaean): swg, cwg, nwg 17, 17, 26, 27, 80 skaergaard (eocene): seg 83 iron ore exploration 78 iron province, melville bugt: nwg 81 island arcs 17, 19, 80 isua: swg 79 isua supracrustal sequence (eoarchaean): swg 17, 17, 18, 80 isukasia: swg 14, 16, 17, 18, 79, 80 isukasia terrane: swg 16 itilli: cwg 79 itilli fm (cretaceous): cwg 57, 58 itilli oil type 89, 90 itsaq gneiss complex (eoarchaean): swg 18 ivigtut cryolite deposit (gardar): swg 78, 79, 80, 82 ivisaartoq: cwg 79 ivittuut: swg 79 ivittuut block: swg 16 j jameson land basin: ceg 55, 55, 74, 79, 85 petroleum potential 85, 86 jan mayen fracture zone: neg 66, 67, 69, 74 jan mayen microcontinent (norway–iceland) 66, 67, 69 josephine headland fm (meso–neoroterozoic): nwg 40 julianehåb batholith (palaeoproterozoic): sg 15, 16, 28, 29, 30, 78 julianehåb granite: sg 29 jurassic, sediments cng, eng 53, 54, 54 neg, ceg 55, 55, 56, 56, 57, 74, 74, 86, 87, 91 swg 76 jyske ås fm (neoproterozoic): cng, eng 42 k kakortokite (gardar): sg 35 kangaamiut basin: swg 85 kangâmiut dykes (palaeoproterozoic): swg 22, 36, 37 kangâmiut-1 exploration well: swg 76, 88, 89 kangeq sequence (cretaceous): swg 76, 76, 89, 90 kangerlussuaq gp (cretaceous–palaeogene: seg 60 kangerdlugssuaq intrusion (palaeogene): seg 62 kangerluk structure: cwg 89 kangerluluk: seg 79 kangerlussuaq (airport): swg 4, 16 kangerlussuaq basin: seg 79 kangersik kiatteq: ceg 2, 31, 38, 39 kangilia fm (cretaceous–paleocene): cwg 57, 58, 59 kanumas 66, 72, 73, 76, 87, 90 kanumas preference areas 87 kap alexander: nwg 40 kap alexander fm (mesoproterozoic): nwg 40 kap bernhard fm (neoproterozoic): cng, eng 42 kap dalton: seg 62 kap dalton gp (eocene): seg 62 kap edvard holm intrusion (palaeogene): seg 62, 79 kap graah gp (devonian): neg 53, 53 kap gustav holm intrusion (palaeogene): seg 62 kap holbæk fm (cambrian): ng 41, 41 kap københavn fm (plio-pleistocene): cng 62, 63 kap kolthoff gp (devonian): neg 53, 53 kap parry intrusion (palaeogene): neg 62 kap simpson intrusion (palaeogene): neg 62 kap stewart gp (triassic–jurassic): ceg 56, 87 kap trautwine fm (meso–neoproterozoic): nwg 40 kap washington: cng 53 kap washington gp (paleocene): cng 48, 53, 54 kap york: nwg 40 kap york meta-igneous complex (neoarchaean): nwg 27 kapisilik terrane: swg 16 karrat: cwg 79 karrat gp (palaeoproterozoic): cwg, nwg 14, 15, 25–27, 27, 28 karrat isfjord: cwg 27 karstryggen: ceg 79 kennedy kanal (channel): nwg 49, 69, 72 ketilidian front: swg 16 ketilidian orogen: seg, sg 14, 14–16, 21, 22, 24, 28–30, 30 gold mineralisation: sg 78, 81 kialineq intrusion (palaeogene): seg 62 kigarsima nappe: cwg 28 kimberlite: swg 36, 37, 44 diamond exploration: swg 78, 80, 82 kimmeridgian marine mudstones: eg 86, 87 kirkespirdalen gold deposit (nalunaq): sg 78, 81 kitsissut sequence (cretaceous), offshore: swg 75, 76, 89 kitsissut (carey islands): nwg 77 kivioq basin: nwg 77, 85 kivioq ridge: nwg 77 kløftelv fm (cambrian): neg 50 knipovich ridge (svalbard) 69, 70 kobberminebugt: sg 79 koch, lauge 11 kolbeinsey ridge (iceland) 66, 69 koldewey platform: neg 74 kome fm (cretaceous): cwg 57, 58 kong oscar fjord gr: neg 50, 51 kort & matrikelstyrelsen (kms: national survey and cadastre) 13 kronprins christian land: eng 34, 35, 40, 41, 45, 46, 47, 51, 122 53, 74 ‘kronprins christian land orogeny’: eng 55 krummedal supracrustal sequence (mesoproterozoic): ceg, neg 2, 38, 38, 39, 39, 41, 50, 51 kvanefjord block: swg 16 l labrador sea 69, 70–72, 76, 85, 88, 90 ladegårdsåen fm (jurassic): cng 54, 54 lady franklin basin: swg 76, 85 lamproites: cwg 36, 37 lamprophyres: swg, swg 37, 44, 82 langø: nwg 79 laurentia 14, 50 laurentian shield 9, 14 lead: ng, cwg, neg 26, 78, 79 lead-zinc deposits citronen fjord: cng 78, 79, 82 maarmorilik (sorte engel): cwg 79, 80, 81 mestersvig: neg 79, 83 legend concept 11–13 licensing rounds, hydrocarbon exploration 87, 88 lilloise intrusion (palaeogene): seg 62 lincoln sea (hav): ng 69, 73, 85 lincoln sea basin: ng 73, 77, 85 lindemans bugt fm (jurassic–cretaceous): neg 56 lithium, prospect/deposit: sg 82 liverpool land basin: ceg 74, 75, 75, 85 petroleum potential 87 llandovery (silurian): ng 47 lomonosov ridge: ng 69, 73, 77 ludlow (silurian): ng 47 lujavrite (gardar): sg 35 lyell land gp (proterozoic): neg 42, 43, 44 m m’clintock orogeny (ordovician), canada maarmorilik: cwg 79 lead-zinc deposit 80, 81 maarmorilik formation: cwg 27 maastrichtian (cretaceous): cwg 57, 58, 59 magnetic data 24, 24, 66, 67, 69–72, 74, 76, 84, 88, 90 magnetochrons 70, 71 majuagaa: swg 79 makkovik orogen (canada) 15 maligât fm (paleocene): cwg 58, 60 mallemuk mountain gp (carb.–permian): cng, eng 53, 54, 54 malmbjerg: ceg 79 molybdenum deposit: ceg 80, 83 maniitsoq block: swg 16, 21 maniitsoq high: swg 76 marinoan (c. 635 ma; tillite gr) 41–43 markham fold belt (ordovician), canada 49 mârmorilik fm (proterozoic): cwg 26, 78 mass balance of inland ice 64, 65 ‘md’ dykes (palaeoproterozoic): swg, swg 37 melville bugt: nwg 4, 76, 79, 85 melville bugt fault: nwg 77 melville bugt graben: nwg 76, 77, 77, 85 melville bugt ridge: nwg 77 merqujôq fm (silurian): ng 48 mesoarchaean supracrustal rocks 19 gneisses 19 intrusive rocks 19, 21 mesoproterozoic rocks 33–35, 35, 38, 39, 40, 41, 44, 49, 50, 79 mineral deposits 80, 82 mesozoic sedimentary onshore basins eg 55, 55, 56, 56, 60 ng 53, 54, 54, 55 wg 57, 58, 59, 59 mesozoic sediments, offshore basins: eg, wg 75, 76, 87 mestersvig: ceg 4, 79 lead-zinc deposit 78, 79, 80 midsommersø dolerites (mesoproterozoic): cng, eng 33, 34, 37, 40 milne land: ceg 79 mineral deposits chromium 79, 80 coal 57, 58, 79, 80, 83 cryolite 78, 79, 80, 82 diamonds 44, 78, 79, 80, 82 gold 78, 79, 80–83, 81 iron 17, 17, 27, 78, 79, 80, 81, 83 lead 26, 78, 79, 80–83 molybdenum 79, 80, 83 pge (platinum group elements) 78, 79, 83 platinum 78, 79, 83 silver 79, 81, 82 uranium 79, 82 zinc 78, 79, 80–83 mines 79, 80–83, 81 miocene (tertiary) glaciation 63 intrusion 83 offshore: ceg, wg 75, 90 volcanic rocks 61, 68, 69 mohns ridge 69, 70 moho 72 molloy fracture zone (svalbard) 69 molybdenum: seg, ceg 78, 79, 83 molybdenum deposit (malmbjerg): ceg 79, 80, 83 morænesø fm (neoproterozoic): cng 41, 41 moriusaq: nwg 79 morris bugt gp (ordovician–silurian): cng, wng 45, 46 morris jesup rise: cng 69, 70 n nagssugtôq subbasin: swg 89 nagssugtoqidian front: swg 16 nagssugtoqidian orogen (palaeoproterozoic): cwg, swg, seg 14, 21, 22–25, 23, 24, 25, 36 mineral deposits 80, 82 nalunaq: sg 79 123 gold mine (kirkespirdalen) 78, 81 nansen fjord: ceg 60 nares land: cng 46 nares stræde (nares strait): nwg, wng 4, 66, 69 nares strait geological structure: nwg, wng 72, 73 nares strait gp (meso–neoproterozoic): nwg 37, 40, 40 narsaq: sg 35, 82 narssârssuk gp (neoproterozoic): nwg 40, 40 narwhale sound fm (ordovician): neg 50 naternaq: cwg 79 nathorst land gp (neoproterozoic): neg 42, 43 national survey and cadastre (kms) 13 nationale geologiske undersøgelser for danmark og grønland, de 2, 4 naujaite (gardar): sg 35 navarana fjord: cng 79 neill klinter gp (jurassic): ceg 56 neoarchaean 15 supracrustal rocks 19, 81 gneisses 19, 20, 81 intrusive rocks 20, 21, 36 neogene (tertiary) 73, 74, 75, 77, 86, 87, 90 neoproterozoic sedimentary onshore basins nwg 39, 40 eg 41, 43 ng 40, 41 orogeny (grenvillian) 38, 39, 50, 51 intrusions 82 neria terrane: swg 16 nertiit kangersivat: ceg 61 nesmith beds (canada) 45 net-veined intrusions (proterozoic): sg 29 new quebec orogen (canada) 15 niaqornaarsuk: sg 79 nickel, prospect: cwg, swg, sg, seg 79 niggli spids thrust sheet: ceg, neg 39, 51 niobium, prospect/deposit: swg, sg 79, 80, 82 nordre isortoq steep belt: swg 15 nordvestfjord: ceg 2, 31, 38, 39 north atlantic craton 15, 16, 26, 27 north atlantic d surveys (nad project) 66 north ungava basin (canada) 89 north water bay: nwg 69, 77 northumberland fm (meso–neoproterozoic): nwg 40 nualik intrusion (palaeogene): seg 62 nûkavsak fm (palaeoproterozoic): cwg 26, 27 nukik-1 & 2 exploration wells: swg 88 nunaoil a/s 66 nuuk basin: swg 76, 85 nuuk: swg 4, 14, 16, 85, 87 nuup kangerlua (godthåbsfjord): swg 16 nuussuaq basin: cwg 57, 58, 59, 59, 76, 79, 89 core wells 86 offshore 76, 89, 90 nuussuaq group (cretaceous–palaeogene): cwg 57 o ocean drilling program (odp) 66, 67, 75 ocean floor ridges 66, 67, 69, 70, 73, 77 oceanic crust 13, 66, 67, 68, 69, 69, 70–73, 74, 75, 77, 87, 88 ocean-to-continent transition zone 13, 66, 67, 68, 69, 77 odp see ocean drilling program offshore 66–77 geology 68, 85 sedimentary basins 85 oil (petroleum potential) 84–91 oil seeps 84, 86, 89, 90 oil showings 86, 89, 90 oligocene hiatus offshore: wg 76 olivine deposit, niaqunngunaq/fiskefjord: swg 79, 82 olympen fm (jurassic): ceg 56 onshore sedimentary basins, see sedimentary basins, onshore ‘open-door’ areas, hydrocarbon exploration 87, 88 ordovician carbonate sediments: ng 41, 45, 46, 47, 49 clastic sediments: ng 46, 48 deep-water trough: ng 46, 48 carbonate sediments: neg 50, 50, 51 ‘fossilik’: swg, see legend explanation [9] sediments: neg 71, 75, 110 offshore: wg 71, 75 sediments with mineral deposits 78, 82 p paamiut basin: swg 85 paamiut terrane: swg 16 pacific fauna 50 palaeogene 12, 57–62, 86 sediments 57, 58, 60, 62, 74, 75, 76, 84, 86, 89 volcanic rocks 61, 61, 68, 69, 71, 74, 74, 79 intrusions 62, 62, 67, 78, 79 mineral deposits 83 palaeoproterozoic unfolded sediments 33, 34, 35, 41 supracrustal rocks 23, 26, 28, 30 intrusive rocks 24, 27, 29 palaeoproterozoic orogenic belts 79 basement in caledonian fold belt: ceg, neg 31, 32, 52 inglefield orogenic belt: nwg 27, 28 ketilidian orogen: sg, seg 28–30, 81 nagssugtoqidian orogen: cwg, swg, seg 22–24 nagssugtoqidian in ammassalik region: seg, seg 24 rinkian fold belt: cwg, nwg 25–27, 78 palaeozoic fold belts: ng, neg 49, 50, 51 sedimentary basins, see sedimentary basins paleocene sediments eng 55 cwg 58, 86, 89, 90 offshore: wg, neg 74, 75, 76, 76 seg 60 paleocene volcanic rocks: cng 54, 71 124 intrusions: seg 62 oceanic crust 68, 69, 72 palladium, prospect/deposit: seg, cwg 79, 83 palnatokes bjerg fm (jurassic): neg 56 pandora havn fm (meso–neoproterozoic): nwg 40 paradisfjeld gp (cambrian): cng 45, 46, 48 passive margins: eg, wg, ng 27, 45, 70 peary land gp (silurian): ng 46, 47, 48 peary land: cng 4, 33, 46, 48, 53 pearya (canada) 48, 49 peralkaline rocks ilímaussaq intrusion (gardar): sw 35, 79 kap washington gp (paleocene): cng 54 peridotite 23, 70, 71 permian sediments ceg, neg 55, 55, 56, 56, 57, 73, 74, 86, 87 cng, eng 53, 54, 54, 57, 86 petermann halvø fm (silurian): ng 45, 46 petroleum appraisal (by us geological survey) 84, 87 petroleum exploration licences 84, 87, 88 petroleum potential 85 ng 84 neg 84–88 wg 86, 88–91 phosphor: cwg 79 picrite (palaeogene): cwg, ceg 60, 61 pillow lavas archaean cwg, disko bugt 26 swg, isua 17, 19 palaeoproterozoic neg 43 cwg, aasiaat/egedesminde 23 swg 29 mesoproterozoic ng, zig-zag dal 35 pingo dal fm (triassic): ceg 56 pituffik (thule air base): nwg 4 place names 13, 69, 78, 79, 113 platinum, prospect/deposit: cwg, swg, seg, seg 78, 79, 83 plays (petroleum) 86 pleistocene sediments (quaternary): cng 62, 63, 74, 75, 87 pliocene sediments (neogene): cng 62, 63, 68, 69, 75, 90 polkorridoren gp (cambrian): cng 45, 46, 48 portfjeld fm (cambrian): cng, eng 41, 42, 45, 46, 49 precambrian greenland shield 12, 21, 31, 33, 39, 54, 84 pridoli (silurian): wng 48 prospecting activities 9, 78, 82 proterozoic sediments in caledonian fold belt neoproterozoic: ceg, neg 41, 42 mesoproterozoic: ceg, eng, neg 33, 38, 39 proterozoic unfolded units neoproterozoic sedimentary basin, thule basin: nwg 39 sedimentary basin, hagen fjord basin: ng 40 mesoproterozoic basalts: ng, sg 33, 34, 35 gardar province: sg 35, 38 sedimentary basins: ng, neg, sg, nwg 33, 35, 38, 39 prøven igneous complex (palaeoproterozoic): nwg 14, 15, 26, 27 pseudo-escarpment (pse): eg 67, 75 pulaskite (gardar): sg 35 pull-apart basin: eng 53, 54, 77 q qaanaaq fm (meso–neoproterozoic): nwg 40 qaqqaarsuk: swg 79 qaqarssuk carbonatite complex (jurassic): swg 44, 82 qarliit taserssuat assemblage: swg 16, 19 qeqertarssuaq fm (palaeoproterozoic): cwg 26, 27 qeqertarsuup tunua (disko bugt): cwg 4, 14, 15 qôrqut granite complex (neoarchaean): swg 16, 21 quaternary 63 quikavsak fm (paleocene): cwg 57, 58, 59, 59 qulleq-1 exploration well: cwg 68, 69, 71, 76, 88 qullissat: cwg 79 qullissat coal deposit (cretaceous): cwg 80, 83 qussuk: swg 79 r rae craton (canada) 15, 26, 27 rapakivi ’granite’ (palaeoproterozoic): seg, sg 28, 30, 30, 38 rare-earth element, prospect/deposit: swg, sg, ceg 79, 80, 82 raukelv fm (jurassic): ceg 56 ravnefjeld formation (permian): ceg 87 rensselaer bay fm (meso–neoproterozoic): nwg 40 reservoir rocks 84, 86, 88–91 reykjanes ridge (iceland) 67, 69 rift basins eg 55, 86 wg 73, 75, 90 rinkian fold belt (proterozoic): cwg, nwg 14, 14, 21, 25, 26, 28 rivieradal gp (neoproterozoic): eng 41, 41, 51 robertson fjord fm (meso–neoproterozoic): nwg 40 ruby, prospect/deposit: swg 79, 80, 81 ryder gletscher gp (cambro-ordovician): ng 45, 46 s saalian (quaternary) 63 salt formation (see also evaporates) 73 samuelsen høj fm (silurian): cng, eng 47 sangamonian (quaternary) 63 saqqaq: cwg 79 sarfartoq: swg 79 sarfartoq carbonatite complex (neoproterozoic): swg 44, 79, 82 sea-floor spreading 60, 66, 67, 68, 69–73, 69, 77 seaward-dipping reflectors 66, 67, 68, 69, 75 sedex-type sulphide deposit: cng 82 sedimentary basins offshore blosseville kyst basin: ceg, seg 75, 85, 88 kivioq basin: nwg 77, 85 lady franklin basin: swg 76, 85 liverpool land basin: ceg 74, 75, 75, 85, 87 onshore franklinian basin: ng 45–48, 84 jameson land basin: ceg 55, 55, 56, 56, 74, 79, 86, 87 kangerlussuaq basin: seg 60 nuussuaq basin: cwg 57, 58, 59, 76, 79, 89, 90 palaeozoic, kong oscar fjord gp: eg 50, 51 wandel sea basin: cng, eng 48, 53, 54, 54, 55, 74, 84, 85 wollaston forland: neg 55, 56 petroleum potential 84–90 seismic data offshore 66, 67, 69, 70, 73–76, 87, 88, 90 onshore 86 seqi: swg 79 seqi olivine mine, niaqunngunaq/fiskefjord: swg 79, 82 sermiligaarsuk terrane: swg 16 sermilik block: swg 16 sermilik: seg 25 sillisissanguit: swg 79 sills, see dykes and sills silurian: ng carbonate sediments 45, 46, 47, 49, 51, 84, 85 turbidite trough 45–48, 46, 48, 51, 78, 85 silver, prospect/deposit: cwg, seg 79, 82 sinarsuk: swg 79 singertât complex (neoarchaean): seg 21 sioraq terrane: swg 16 sisimiut basin: swg 85, 89 sisimiut intrusive complex (palaeoproterozoic): swg 24 skærgaard: seg 62, 79 skaergaard intrusion (eocene): seg 62, 62, 78, 79, 83 gold and platinum deposit: seg 79, 83 skagen gp (cambrian): ng 45, 48 skjoldungen intrusions (neoarchaean): seg 21 skolithos: neg 43 slibestensfjeldet fm: cwg 57, 58 slottet fm (cambrian): neg 43 smallefjord sequence (neoproterozoic): neg 39 smith sound: nwg 69, 72 smith sound gp (meso–neoproterozoic): nwg 37, 40, 40 sodalite foyaite (gardar): sg 35 soft-bodied fossil fauna: cng 45 sonntag bugt fm (meso–neoproterozoic): nwg 40 sortebakker fm (carboniferous): eng 53, 54 sortis gp (palaeoproterozoic): swg 29, 29 source rocks (oil and gas) 84 offshore 87–91 onshore 84, 86 spitsbergen fracture zone (greenland–svalbard) 69, 70 spreading axis 67, 68, 69, 71–73, 85 spreading ridge 67, 70 stendalen: sg 79 storø: swg 79 structural trend lines 13 sunrise pynt straight belt (palaeoproterozoic): nwg 28 svalbard (norway) 38, 69, 70, 85 svartenhuk formation (paleocene): cwg 60 sverdrup basin (arctic canada) 77, 90, 91 t taartoq: swg 16, 79 tantalum, prospect/deposit: swg, sg 79, 80, 82 tarssartôq dykes (palaeoarchaean): swg 37 târtoq gp (mesoarchaean): swg 16, 19 taserssuaq tonalite (mesoarchaean): swg 16, 21 tasiilaq: seg 79 tasiusarssuaq terrane (archaean): swg 16 tavsen iskappe gp (cambro-ordovician): cng, wng 45, 46 ‘td’ dykes (cretaceous): swg, swg 37 tectono-metamorphic provinces 12 terranes archaean 15, 16, 17, 19–21 palaeozoic 48, 49 tertiary, mainly palaeogene 60–63 basalts cwg 60 eg 61, 62 offshore: eg 74, 85 offshore: wg 76, 85 dykes and sills: eg 37, 62 intrusions: eg 62 sediments onshore cwg 57, 58, 59 eng 53, 54, 55, 62 seg 60 offshore eg 73–75 wg 75–77 volcanics: eg, wg 85 thetis basin: neg 74, 85 thorium, prospect/deposit: sg 79, 82 thule air base (pituffik) 4 thule sgp (meso–neoproterozoic): nwg 39, 40, 40, 77 thyra ø fm (palaeogene): eng 54, 55 tikiusaaq carbonatite complex (jurassic): swg 44, 82 tillite gp (neoproterozoic): ceg, neg 41–43, 43, 51 tillitic rocks cng 41 neg 42, 43 titanium, prospect/deposit: 79, 83 topographic base 13 toqqusap nunaa: swg 22 trail ø group (carboniferous): neg 53 transform faults 13, 67, 68, 69, 70, 72, 76 transition zone, offshore 13, 66, 67, 77 transitional crust 67, 68, 69, 69, 70, 71, 73 tre brødre: swg 16 tre brødre terrane: swg 16, 17, 20 triassic sediments ceg 56, 56, 87 cng, eng 53, 54, 54 neg 56, 57 nwg 91 trolle land fault zone: cng, eng 53 trolle land gp (permo-triassic): cng 53, 54, 54 tupertalik carbonatite (mesoarchaean): swg 21, 44 turbidite flysch cambro-silurian: ng 41, 45, 46, 46, 47, 48, 48 125 palaeoproterozoic: swg, cwg 23, 26 cretaceous: cwg, seg 57, 58, 59, 60 palaeogene: cwg 89, 90 u ubekendt ejland: cwg 60 umîvik dykes (palaeoproterozoic): seg, seg 37 ungava fracture zone 69, 71, 76, 89, 90 upernavik: nwg 4, 14, 85 upernivik ø: cwg 28 uranium, prospect/deposit: sg 79, 82 uranium concentrations in stream sediments: wg 21 utm projection 13 uummannaq: cwg 14 v vaigat fm (paleocene): cwg 58, 59 vallen gp (palaeoproterozoic): swg 28, 29 vanadium, prospect/deposit: swg, seg 79, 83 vandredalen thrust sheet: eng 40, 41 vardekløft fm (jurassic): ceg 56 victoria fjord: cng, wng 14, 14, 48 vilddal gp (devonian): neg 52, 53 vølvedal gp (cambro-ordovician): cng, wng 46, 46, 48 w wandel dal: cng 41 wandel sea basin: cng, eng 48, 53, 54, 54, 55, 74, 84, 85 wandel sea strike-slip mobile belt: ng 53 wandel valley fm (ordovician): ng 41 washington land: wng 45, 46, 79, 82 washington land gp (silurian): ng 47 weichselian (quaternary) 63, 65 wells, exploration, offshore 66, 68, 69, 71, 73, 75, 76, 84, 87, 88, 89, 89 wells, exploration, onshore 58, 59, 86, 87 wenlockian (silurian): ng 47 werner bjerge complex (tertiary): ceg 62 windows in caledonian orogen: neg, cng 31, 43, 50, 51 wisconsinan (quaternary) 63 wolfram, prospect: swg, ceg, neg 79 wollaston forland (jurassic–cretaceous): neg 55, 56 wolstenholme fm (meso–neoproterozoic): nwg 40 wordie creek fm (triassic): ceg 56 wulff land: wng 46 y yermak plateau (svalbard) 69, 70 ymer ø: neg 79 ymer ø gp (neoproterozoic): neg 42, 43, 44 ymers gletscher fm (silurian): cng 45, 46 yttrium, prospect 82 z zig-zag dal basalt fm (mesoproterozoic): cng, eng 33, 35, 40, 41, 51 zinc, prospect/deposit: ng, cwg, neg 26, 78, 79, 80–83 zirconium, prospect/deposit: sg, ceg 79, 80, 82 126 geological survey of denmark and greenland bulletin 16, 2008, 1-66 geological survey of denmark and greenland bulletin 16 · 2008 evaluation of the quality, thermal maturity and distribution of potential source rocks in the danish part of the norwegian–danish basin henrik i. petersen, lars h. nielsen, jørgen a. bojesen-koefoed, anders mathiesen, lars kristensen and finn dalhoff geological survey of denmark and greenland ministry of climate and energy geological survey of denmark and greenland bulletin 16 keywords norwegian–danish basin, source rocks, generation potential, maturation cover photomicrograph of alginite-rich organic matter from lacustrine sediments of the frederikshavn formation in the terne-1 well (200–210 m depth; see also fig. 39); the photograph is taken in fluorescence-inducing blue light. the sub-circular structure (c. 200 µm across) is an algal body that resembles the freshwater green alga botryococcus. chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, geological institute, university of aarhus; minik rosing, geological museum, university of copenhagen; finn surlyk, department of geography and geology, university of copenhagen scientific editor of this volume: jon r. ineson editorial secretaries: jane holst and esben w. glendal referees: anthony m. spencer (norway) and an anonymous referee illustrations: stefan sølberg, with contributions from jette halskov digital photographic work: benny m. schark layout and graphic production: annabeth andersen printers: schultz grafisk, albertslund, denmark manuscript received: 7 december 2007 final version approved: 25 august 2008 printed: 28 november 2008 issn 1604-8156 isbn 978-87-7871-232-5 geological survey of denmark and greenland bulletin the series geological survey of denmark and greenland bulletin replaces geology of denmark survey bulletin and geology of greenland survey bulletin. citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 16, 66 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk or at www.geus.dk/publications/bull © de nationale geologiske undersøgelser for danmark og grønland (geus), 2008 for the full text of the geus copyright clause, please refer to www.geus.dk/publications/bull 3 contents abstract . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5 introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7 aims of this study . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7 tectonic setting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10 depositional development and stratigraphy . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 14 zechstein evaporites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 14 early–middle triassic clastic deposition . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 14 late triassic clastic and evaporitic deposition and marine flooding . . . . . . . . . . . . . . . . . . . . . 16 hettangian – early pliensbachian transgression and basin expansion . . . . . . . . . . . . . . . . . . . . 16 late pliensbachian – early aalenian sea-level fluctuations . . . . . . . . . . . . . . . . . . . . . . . . . . . . 22 late early – middle jurassic uplift and erosion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 22 late middle – late jurassic basin expansion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 24 cretaceous continued basin expansion and chalk deposition . . . . . . . . . . . . . . . . . . . . . . . . . . 26 cenozoic clastic deposition and neogene exhumation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 27 regional coalification curves corrected for net exhumation . . . . . . . . . . . . . . . . . . . . . . . . . . . . 28 shale velocity-based curve . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 29 chalk velocity-based curve . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 33 depth to the top of the oil window . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 33 distribution and thermal maturity of potential source rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . 36 pre-permian units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 36 permian units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 37 triassic units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 37 lower jurassic units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 38 f-i member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 39 f-ii member . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 39 f-iii and f-iv members . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 40 middle jurassic units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 49 upper jurassic – lower cretaceous units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 52 lower cretaceous units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 56 potential reservoirs . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 56 gassum reservoir . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 56 haldager sand reservoir . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57 additional reservoirs . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57 discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 58 source-rock quality and distribution . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 58 source-rock maturity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 59 reservoirs and migration . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 61 active mesozoic petroleum system? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 62 conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 62 acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 62 references . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 63 4 5 abstract authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: hip@geus.dk *present address: vattenfall a/s, støberigade 14, dk-2450 copenhagen sv, denmark. petersen, h.i., nielsen, l.h., bojesen-koefoed, j.a., mathiesen, a., kristensen, l. & dalhoff, f.* 2008: evaluation of the quality, thermal maturity and distribution of potential source rocks in the danish part of the norwegian–danish basin. geological survey of denmark and greenland bulletin 16, 66 pp. the quality, thermal maturity and distribution of potential source rocks within the palaeozoic–mesozoic succession of the danish part of the norwegian–danish basin have been evaluated on the basis of screening data from over 4000 samples from the pre-upper cretaceous succession in 33 wells. the lower palaeozoic in the basin is overmature and the upper cretaceous – cenozoic strata have no petroleum generation potential, but the toarcian marine shales of the lower jurassic fjerritslev formation (f-iii, f-iv members) and the uppermost jurassic – lowermost cretaceous shales of the frederikshavn formation may qualify as potential source rocks in parts of the basin. neither of these potential source rocks has a basinwide distribution; the present occurrence of the lower jurassic shales was primarily determined by regional early middle jurassic uplift and erosion. the generation potential of these source rocks is highly variable. the f-iii and f-iv members show significant lateral changes in gene ration capacity, the best-developed source rocks occurring in the basin centre. the combined f-iii and f-iv members in the haldager-1, kvols-1 and rønde-1 wells contain ‘net source-rock’ thicknesses (cumulative thickness of intervals with hydrogen index (hi) >200 mg hc/g toc) of 40 m, 83 m, and 92 m, respectively, displaying average hi values of 294, 369 and 404 mg hc/g toc. the mors-1 well contains 123 m of ‘net source rock’ with an average hi of 221 mg hc/g toc. parts of the frederikshavn formation possess a petroleum generation potential in the hyllebjerg-1, skagen-2, voldum-1 and terne-1 wells, the latter well containing a c.160 m thick highly oil-prone interval with an average hi of 478 mg hc/g toc and maximum hi values >500 mg hc/g toc. the source-rock evaluation suggests that a mesozoic petroleum system is the most likely in the study area. two primary plays are possible: (1) the upper triassic – lowermost jurassic gassum play, and (2) the middle jurassic haldager sand play. potential trap structures are widely distributed in the basin, most commonly associated with the flanks of salt diapirs. the plays rely on charge from the lower jurassic (toarcian) or uppermost jurassic – lowermost cretaceous shales. both plays have been tested with negative results, however, and failure is typically attributed to insufficient maturation (burial depth) of the source rocks. this maturation question has been investigated by analysis of vitrinite reflectance data from the study area, corrected for post-early cretaceous uplift. a likely depth to the top of the oil window (vitrinite reflectance = 0.6%ro) is c. 3050–3100 m based on regional coalification curves. the frederikshavn formation had not been buried to this depth prior to post-early cretaceous exhumation, and the potential source rocks of the formation are thermally immature in terms of hydrocarbon generation. the potential source rocks of the fjerritslev formation are generally immature to very early mature. mature source rocks in the danish part of the norwegian–danish basin are thus dependent on local, deeper burial to reach the required thermal maturity for oil generation. such potential kitchen areas with mature fjerritslev formation source rocks may occur in the central part of the study area (central–northern jylland), and a few places offshore. these inferred petroleum kitchens are areally restricted, mainly associated with salt structures and local grabens (such as the fjerritslev trough and the himmerland graben). 6 4º e 58 ºn 57 ºn 56 ºn 55 ºn 6º e 8º e 10 ºe 12 ºe 14 ºe g er m an y horn graben sw ed en sk ån e jy lla nd sj æ lla nd sk ag er ra k k at te ga t b al tic s ea n or w ay te rn e1 a nh ol t1, 4 fl yv bj er g1 sk iv e1 ib en ho lt1 d -1 k -1 d an is h c en tr al g ra be n 10 /7 -1 10 /8 -1 11 /1 01 11 /9 -1 sl ag el se -1 ve m b1 v in di ng -1 r -1 g la m sb je rg -1 g ri nd st ed -1 je lli ng -1 fi g. 6a fi g. 6b fi g. 4 fi g. 5a fi g. 5b fr ed er ik sh av n1 , 2 , 3 vo ld um -1 r øn de -1 g as su m -1 h ob ro -1 m ej ru p -1 r ød di ng -1 fa rs ø1 å rs -1 h yl le bj er g1 ve ds te d1 sæ by -1 n øv lin g1 m or s1 k vo ls -1 b ør gl um -1 h al da ge r1 la vø -1 h an s1 fj er ri ts le v2 f1 in ez -1 d -1 l1 t1 13 /2 -u -2 h or se ns -1 m ar gr et he ho lm -1 st en lil le w el ls ø re su nd -5 ,-7 u lle rs le v1 r -1 c -1 j1 sk ag en -2 w el ls g eo se ct io ns fa ul t n at io na l b or de r h ig h 10 0 km so rg en fre i– to rn qu ist z on e bø rg lu m f au lt f je rr its lev f au lt n or w eg ia n– d an ish b as in ri ng kø bi ng –f yn h ig h sk ag er rak –k att eg at pla tfo rm h i. g r. fj er rit sle v tr ou gh ea st n or th se a h ig h tt z fa rs un d ba si n h ald ag er fa ul t fe lic ia -1 , 1 a st av an ge r pl at fo rm li st a fa ul t bl oc k c om pl ex eg er su nd basi n fi g. 1 . m ap sh ow in g w el l p os iti on s a nd th e st ru ct ur al o ut lin e of th e n or w eg ia n– d an ish b as in , t he s or ge nf re i– to rn qu ist z on e, in cl ud in g th e fj er ri ts le v tr ou gh a nd th e h im m er la nd g ra be n (h i. g r.) , an d th e sk ag er ra k– k at te ga t p la tf or m . t he m aj or fa ul t s ys te m s i nd ic at ed a re th os e ac tiv e in th e m es oz oi c (c f. fi gs 2 , 3 ). po sit io ns o f g eo se ct io ns d isp la ye d in f ig s 4 –6 a re a lso sh ow n. n am ed w el ls ar e di sc us se d in th e te xt o r u se d in lo gpa ne ls; u nn am ed w el ls ar e in di ca te d to il lu st ra te th e de ep w el ls us ed in m ap c om pi la tio n (e .g . f ig . 1 5) . a ss es sm en t o f t he so ur ce -r oc k po te nt ia l i s b as ed o n da ta fr om w el ls in di ca te d in it al ic s. t t z , t ei ss ey re –t or nq ui st z on e. 7 this study focuses on an evaluation of source rocks in the danish part of the norwegian–danish basin, which is large ly equivalent to the danish embayment and the danish subbasin of earlier workers (figs 1, 2; sorgenfrei & buch 1964; larsen 1966; michelsen 1975, 1978, 1989a, b; bertelsen 1978). the danish part of the norwegian–danish basin (referred to hereafter as ‘the study area’) covers onshore denmark and danish offshore territory extending as far west as the eastern margin of the danish central graben in the north sea. several sub-basins occur in the area, including the himmerland graben and the fjerritslev trough (fig. 2). since 1935, more than 60 deep wells have been drilled throughout the study area (see fig. 2), yielding valuable in formation related to petroleum prospectivity (fig. 1). some of the wells were drilled for geothermal energy or gas storage, but the majority of the wells in the basin were hydrocarbon exploration wells, the main target being the mesozoic succession (sorgenfrei & buch 1964; nielsen & japsen 1991). the mesozoic petroleum system that relies on lower jurassic source rocks of the fjerritslev formation is considered to be the principal petroleum system in the norwegian–danish basin. in the study area, the dominant play models have involved sandstones of the upper triassic – lowermost jurassic gassum formation or the middle jurassic haldager sand and bryne formations charged from source rocks within the lower jurassic (fjerritslev formation) or the uppermost jurassic – lowermost creta ceous (børglum, frederikshavn, tau and mandal forma tions). these plays have been tested by more than 30 wells placed on structures or stratigraphic pinch-outs. the results of wildcat drilling in the study area have so far been disappointing as only poor indications of hydrocarbons have been encountered in the wells. the primary source rocks, the f-iii and f-iv members of the fjerritslev formation, have previously been determined to be immature to very early mature in the wells drilled. the wells have mainly been drilled on positive structures, but regional seismic data also support a relatively shallow burial depth for the lower jurassic source-rock intervals over much of the study area. sufficient burial of the source rocks prior to late cretaceous – palaeogene inversion events and neogene exhumation may have occurred in localised depressions, however, such as in rim synclines developed adjacent to salt diapirs. the mesozoic succession has also been tested in the norwegian part of the basin and the border zone by c. 10 wells, also with poor results in terms of hydrocarbons. the norwegian farsund basin (fig. 1), immediately north of the study area, has recently been drilled, and the well encountered jurassic sandstones and lower permian – carboniferous rocks without indications of hydrocarbons. however, oil probably charged from lower jurassic (fjerritslev formation) and upper jurassic (tau formation) source rocks is under production from middle–upper jurassic sandstones (sandnes reservoir) in the yme field, situated in the egersund basin some 75–100 km west of the study area in the north-western part of the nor wegian–danish basin (fig. 1; husmo et al. 2003). aims of this study the primary aim of this bulletin is to provide a comprehensive assessment of the quality, thermal maturity and distribution of potential source rocks within the palaeozoic–mesozoic succession of the central area of the danish part of the norwegian–danish basin (fig. 2). source-rock quality has been assessed by evaluation of screening data (total organic carbon (toc) contents, s2 pyrolysis yields, hydrogen index (hi) values) from more than 4000 samples from the pre-upper cretaceous succession in 33 wells (fig. 1). tmax values and vitrinite reflec tance (vr) data have been used to determine the thermal maturity of potential source rocks. in order to assess the maturity of the various sourcerock levels in the study area, an understanding of neogene exhumation is essential. a previously published coalification curve that was corrected for post-early cretaceous exhumation on the basis of sonic velocity data suggested that the lower jurassic source-rock interval is immature to very early mature and the uppermost jurassic – lowermost cretaceous interval is immature in the danish wells (peter sen et al. 2003a). in this study, new regional coalification (matu ration) curves are presented based on a large number of vr measurements (a total of 560 vr values from 26 wells). the present-day depths of the samples have been corrected for (1) post-early cretaceous net-exhumation magnitudes derived from chalk velocities, and (2) postearly cretaceous net-exhumation magnitudes derived from re-evaluation of the shale sonic ve locity data. the secondary aim of this bulletin is to review the potential mesozoic reservoirs in the study area, another critical element in a viable petroleum system. introduction 8 10 0 km > 9 km fa ul t zo ne bo un da ry st ud y ar ea 8– 9 km 7– 8 km 6– 7 km 3– 4 km 5– 6 km 4– 5 km 2– 3 km 1– 2 km 4º e 58 ºn 57 ºn 56 ºn 55 ºn 6º e 8º e 10 ºe 12 ºe 14 ºe 16 ºe n o r w e g i a n – d a n i s h b a s i n n o r w e g i a n – d a n i s h b a s i n n o r w e g i a n – d a n i s h b a s i n r i n g k ø b i n g – f y n h i g h r i n g k ø b i n g – f y n h i g h r i n g k ø b i n g – f y n h i g h fi g_ 2 e as t n o rt h e as t n o rt h s ea h ig h s ea h ig h e as t n o rt h s ea h ig h f ar su n d b as in f ar su n d b as in f ar su n d b as in eg er su nd eg er su nd b as in b as in eg er su nd b as in r ø n n e r ø n n e g ra b en g ra b en r ø n n e g ra b en fj er ri ts le vt ro ug h fj er ri ts le vt ro ug h fj er ri ts le vt ro ug h s o r g e n f r e i – t o r n q u i s t z o n e s o r g e n f r e i – t o r n q u i s t z o n e s o r g e n f r e i – t o r n q u i s t z o n e h i. h i. g r. g r. h i. g r. fi g. 2 . d ep th s tr uc tu re m ap o f t he to p pr ez ec hs te in s ur fa ce , t he d ee pe st s ur fa ce th at c an b e m ap pe d by c on ve nt io na l s ei sm ic d at a in th e n or w eg ia n– d an ish b as in ; t he e xt en t o f t he s tu dy a re a co ver ed b y th is bu lle tin is a lso sh ow n. n ot e th e la rg e th ic kn es se s i n th e fj er ri ts le v tr ou gh a nd th e h im m er la nd g ra be n (h i. g r.) . m od ifi ed fr om v ej bæ k (1 99 7) . 9 4º e 58 ºn 57 ºn 55 ºn 10 0 km 6º e 8º e 10 ºe 12 ºe 14 ºe 16 ºe ba se m en t be lo w t pz pa la eo zo ic b el ow t pz z ec hs te in s al t pi llo w z ec hs te in s al t di ap ir c al ed on ia n de fo rm at io n fr on t li m it of z ec hs te in d ep os its fa ul t n at io na l b or de r fi g. 3 . m ap sh ow in g th e st ra ta su bc ro pp in g th e to p pr ez ec hs te in su rf ac e (a nd la te ra l c or re la tiv e su rf ac e) a nd th e di st ri bu tio n of k no w n sa lt st ru ct ur es . t he b lu e lin e in di ca te s t he li m it of th e z ec hs te in de po sit s ( ve jb æ k 19 97 ). t pz , t op p re -z ec hs te in . 10 the danish part of the norwegian–danish basin is a wnw–ese-trending intracratonic basin, containing permian–cenozoic strata, that to the south is bounded by elevated precambrian basement of the ringkøbing–fyn high and to the north-east and east by the fennoscandian border zone. the fennoscandian border zone consists of the sorgenfrei–tornquist zone and the skagerrak–kattegat platform that westwards passes into the stavanger platform north-east of the egersund basin; the zone marks the transition to the stable precambrian baltic shield (fig 1; sorgenfrei & buch 1964; bergström 1984; eugeno-s working group 1988; michelsen & nielsen 1991, 1993; vejbæk 1997; nielsen 2003). the ringkøbing–fyn high separates the norwegian–danish basin from the north german basin and was probably formed contemporaneously with the norwegian–danish basin as an area of less crustal stretching (figs 1–3). on the skagerrak–kattegat platform, the mesozoic section is relatively undisturbed; it onlaps tilted fault blocks comprising precambrian crystalline basement, lower palaeozoic and lower permian strata and wedges out towards the north-east (figs 4, 5a). the sorgenfrei–tornquist zone is a highly block-faulted 30–50 km wide zone that runs se–nw from the rønne graben in the baltic sea across southern sweden through the kattegat and northern jylland to the skagerrak, where it turns westwards across the norwegian shelf (figs 1–3, 5b). the zone includes the deep fjerritslev trough with a zechstein–mesozoic succession that locally is more than 9 km thick, and the shallower farsund basin where the zechstein–mesozoic section locally attains a thickness of slightly more than 6 km (figs 2, 6). the principal rifting phase of the norwegian–danish basin and the sorgenfrei–tornquist zone is defined by the occurrence of tilted fault blocks with basement rocks and lower palaeozoic strata unconformably overlain by permian rocks (figs 4–6; liboriussen et al. 1987; vejbæk 1989, 1997; michelsen & nielsen 1991, 1993; jensen & schmidt 1993; christensen & korstgård 1994; vejbæk & britze 1994). the crests of the fault blocks are deeply truncated and this top pre-zechstein surface is the deepest regional surface that can be mapped on reflection seismic data in the norwegian–danish basin and the fennoscandian border zone (fig. 3). the surface is relatively flat and smooth, in dicating pronounced erosion prior to the zechstein transgression. the unconformity is penetrated by wells that testify to the presence of precambrian crystalline rocks on the ringkøbing–fyn high (glams bjerg-1, grindsted-1, ibenholt-1, jelling-1) and the skager rak–kattegat plat form (frederikshavn-1) and lower pa laeozoic sedimentary rocks in the norwegian–danish basin and fennoscandian border zone (figs 1, 3; nøvling-1, rønde-1, slagelse-1, terne-1; sorgenfrei & buch 1964; poulsen 1969, 1974; larsen 1971, 1972; christensen 1971, 1973; nielsen & japsen 1991; michelsen & nielsen 1991, 1993). the shallow iku/sintef borehole 13/2-u-2 encountered silurian strata subcropping the quaternary northeast of the farsund basin (fig. 1). the tilted fault-block crests are deeply truncated by the mid-permian unconformity showing that regional postrift thermal subsidence was somewhat delayed (vejbæk 1997). the unconformity that defines the base of the postrift sequence is overlain by a relatively complete succession of zechstein salts and carbonates, triassic clastics, carbonates and salts, jurassic – lower cretaceous clastics, upper cretaceous chalks and cenozoic clastics that attain a thickness of 5–6.5 km along the basin axis. a similar thickness of the post-rift succession is found in the farsund basin, where the succession is more than 3 sec. (twt) thick in t-1 25 km fig. 4 0.0 1.0 2.0 3.0 4.0 5.0 6.0 sw t w t s ec . danish central graben norwegian–danish basin k-1f-1 tectonic setting 11 places, corresponding to slightly more than 6 km (jensen & schmidt 1993; vejbæk & britze 1994), whereas the succession is more than 9 km thick locally in the fjerritslev trough and the himmerland graben (fig. 2). isochore maps of the triassic and jurassic – lower cretaceous successions show a relatively uniform regional thickness over most of the basin, except for areas influenced by local halokinetic movements, indicating relatively uniform thermal subsidence (vejbæk 1989, 1997; britze & japsen 1991; japsen & langtofte 1991). although the thick upper permian – triassic succession indicates rapid subsidence that exceeds rates normally associated with post-rift thermal contraction (fig. 7), a prolonged or new rifting phase is precluded by the general lack of pronounced extensional faulting in the mesozoic succession; phase transformations in the deep crust have been proposed to explain the rapid, early post-rift subsidence (vejbæk 1989, 1997). deposition and preservation of great thicknesses of mesozoic sediments in the himmerland graben and the fjerritslev trough were facilitated by transtensional strike-slip move ments in the sorgenfrei–tornquist zone or large-scale salt movements (pegrum 1984; vejbæk 1989; christensen & korstgård 1994; mogensen 1994, 1996). growth of salt structures influenced mesozoic deposition locally. the ringkøbing–fyn high was uplifted significantly in early middle jurassic times and the norwegian–danish basin became tilted to the north-east (michelsen 1978; koch 1983; andsbjerg et al. 2001; nielsen 2003). the uplift and tilting caused progressively deeper erosive truncation of the lower jurassic and triassic across the basin towards the ringkøbing–fyn high, where erosion removed the entire lower jurassic and much of the triassic on the most elevated parts of the high (fig. 8). in contrast, subsidence continued during middle jurassic times in the sorgenfrei–tornquist zone, as shown by well sections in the øresund region (øresund-5, -7), in the kattegat area (terne-1, anholt-4), and in the fjerritslev trough (børg lum-1, fjerritslev-2, flyvbjerg-1, j-1, haldager-1, vedsted1), but at a much lower rate than in triassic – early jurassic times (fig. 7). danish well sections on the skagerrak –kattegat platform (frederikshavn-1, -2, -3, skagen-2, sæby-1) indicate that only limited erosion occurred north of the fjerritslev trough. regional subsidence gradually resumed during late middle – late jurassic times and tectonic tranquillity generally prevailed, except for local salt movements, until inversion occurred in the sorgenfrei–tornquist zone. a late cretaceous – palaeogene age is generally accepted for the inversion, which is interpreted to have been caused by compression related to a change in the regional stress field from extensional to compressional, linked to alpine deformation and opening of the north atlantic (liboriussen et al. 1987; ziegler 1990; michelsen & nielsen 1991, 1993; mogensen & korstgård 2003). the inversion may have occurred in several phases, however, and it is assumed to have begun in turonian times or earlier in the south-east, with 1.5–3 km of uplift in the skåne–bornholm area, and propagated north-westwards with decreasing intensity of deformation (berthelsen 1992; mogensen & jensen 1994; michelsen 1997; petersen et al. 2003a; japsen et al. 2007). after cessation of the inversion, a new quiescent tectonic phase began with regional subsidence of the greater north sea basin. contemporaneously with the regional down-warping of southern scandinavia, significant uplift and erosion began to influence parts of the norwegian–danish basin and the ringkøbing–fyn high in neogene times (japsen 1993; jensen & schmidt 1993; japsen et al. 2002a, b, 2007). j-1 ne skagerrak–kattegat platform felicia-1/1a (projected) sæby-1 (projected) norwegian–danish basin sorgenfrei–tornquist zone 0.0 1.0 2.0 3.0 4.0 5.0 6.0 t w t s ec . cenozoic u. cretaceous l. cretaceous u. jurassic l.–m. jurassic triassic zechstein rotliegend fig. 4. regional geosection from the central graben in the south-west to the skagerrak–kattegat platform in the north-east showing pinch-out of the mesozoic strata on the skagerrak–kattegat platform. for location of section, see fig. 1. twt, two-way travel time. 12 nøvling-1 ringkøbing–fyn high norwegian–danish basin haldager-1 flyvbjerg-1 frederikshavn-1 fjerritslev trough sorgenfrei–tornquist zone ne sw skagerrak–kattegat platform 10 km 0 1 2 3 4 t w t ( se c. ) 0 1 2 3 4 t w t ( se c. ) zechstein middle–lower jurassic lower–middle triassic cenozoic upper cretaceous lower cretaceous/ upper jurassic upper triassic rotliegend basement palaeozoic basement lower palaeozoic triassic triassic triassic jurassic jurassic jurassic cambrian cambrian l. cretaceous? ordovician-silurian ordovician-silurian rotliegendeu.car.u.car. u. carboni-ferous børglum fault zechstein 0 1 2 3 t w t ( se c. ) z z r r 5 km ? ? hans-1 l. cretaceous? jurassic triassic zechstein rotliegend u. carboniferous ordovician–silurian cambrian basement sw ne sorgenfrei–tornquist zone a b fig. 5. geosections modified from michelsen & nielsen (1991); for locations, see fig. 1. a: regional geosection across the danish part of the norwegian–danish basin showing thickening of the mesozoic section in the fjerritslev trough / sorgenfrei–tornquist zone. b: geosection from the kattegat through the hans-1 well; note palaeozoic fault blocks. 13 upper cretaceous & danian cenozoic internal reflectors quaternary seabed lower cretaceous upper reflector truncation jurassic middle jurassic (haldager sand fm) lower jurassic triassic triassic salt (oddesund fm) zechstein salt palaeozoic basement s n s n seabed 0 1 2 3 4 5 6 7 0 1 2 3 4 5 t w t ( se c. ) t w t ( se c. ) k-1 j-1 farsund basin fjerritslev fault fjerritslev trough 20 km a b fig. 6. geosections modified from jensen & schmidt (1993); for locations, see fig 1. a: geosection from the norwegian–danish basin across the fjerritslev fault and the farsund basin. the middle jurassic haldager sand reservoir overlying the lower jurassic potential source rocks is marked in red. note palaeozoic fault blocks overlain by zechstein salt. b: geosection from the fjerritslev trough. 14 a review of the permian–cenozoic depositional evolution of the basin is presented here to provide a framework for discussion of potential components of a petroleum system in the danish portion of the norwegian–danish basin. for detailed, comprehensive accounts of the basin evolution and stratigraphy, the reader is referred to michelsen et al. (2003) and nielsen (2003). zechstein evaporites deposition of the post-rift succession was initiated in late permian times with the accumulation of thick zechstein evaporites in most of the norwegian–danish basin. mar ginal facies were developed along parts of the ring købing–fyn high in late permian times, and as the lower triassic bunter shale formation seems to rest on deeply weathered basement in the grindsted-1 well, it is likely that the high formed a barrier between the southern and northern zechstein basins (ziegler 1982; stemmerik et al. 1987; vejbæk 1997). later mobilisation of zechstein salt led to the formation of salt structures such as pillows and diapirs (fig. 3), and the continued growth of salt structures influenced mesozoic deposition in places. towards the east, the evaporites are replaced by thin clastic zechstein deposits as shown by the terne-1 and hans-1 wells in the sorgenfrei–tornquist zone (michelsen & nielsen 1991, 1993). farther north-west in the sorgenfrei–tornquist zone, however, evaporites are present in the fjerritslev trough and the farsund basin (e.g. liboriussen et al. 1987; jensen & schmidt 1993; christensen & korstgård 1994; vejbæk 1997). early– middle triassic clastic deposition in early triassic times, the depositional environment changed to more continental conditions. similarity in depositional facies in the north german and norwegian–danish basins indicates that the two basins were connected, at least periodically (bertelsen 1978; michelsen & clausen 2002). the facies encountered in the jelling-1 and grindsted-1 wells located on the northern flank of the ringkøbing–fyn high belong to the bunter shale, bunter sandstone, ørslev, falster and tønder formations; these formations may be traced farther northwards to the mors-1 d ep th ( m ) 0 500 1000 1500 late triassic early jurassic middle jurassic late jurassic e.cret. time (ma) 220 210 200 190 180 170 160 150 140 ringkøbing–fyn high basin centre skagerrak–kattegat platform himmerland graben fjerritslev trough fig. 7. subsidence curves for the late triassic – early cretaceous constructed for different structural elements/sub-basins within the norwegian–danish basin, skagerrak – kattegat platform and ringkøbing–fyn high based on a sequence stratigraphic break-down (nielsen 2003). depositional development and stratigraphy 15 sw ne ma dolomitic limestones fluvial sandstones alluvial conglomerates, sandstones sabkhas and lacustrine calcareous, evaporitic mudstones estuarine/lagoonal sandstones, heteroliths, mudstones, coal beds shoreface sandstones/siltstones offshore mudstones; occasionally sandy or silty lacustrine mudstones hiatus source rocks norian rhaetian hettangian sinemurian pliensbachian toarcian aalenian bajocian bathonian callovian oxfordian kimmeridgian volgian ryazanian chronostratigraphy sequence key surfaces lithostratigraphy tectonics norwegian–danish basin rkf stz skp fr 3 150.7 142.0 154.1 159.4 164.4 169.2 176.5 180.1 189.6 195.3 201.9 205.7 209.6 fr 2 fr 1 bø 1 fl 1 ha 3 ha 2 ha 1 fj 10 fj 9 fj 8 fj 7 fj 6 fj 5 fj 4 fj 3 fj 2 fj 1 ga 1 vi 1 frederikshavn fm børglum fm flyvbjerg fm haldager sand fm f-iv mb f-iii mb f-ii mb f-ib f-i mb f-ia gassum fm vinding fm sk od fj er ri ts le v fm 27 25 24 23 22 20 19 18 17 16 15 14 13 12 11 9 7 1 26 25 24 23 22 21 20 19 18 17 16 15 14 13 12 11 10 9 5 renewed regional subsidence including ringkøbing– fyn high basin expansion fault-controlled subsidence of sorgenfrei– tornquist zone uplift of ringkøbing– fyn high, ne-tilting of basin thermal subsidence local faulting and halokinesis fjerritslev trough himmerland graben local incision erosion over salt structures condensed section base middle jurassic unconformity ‘mid-cimmerian unconformity’ sb mfs fig. 8. time-stratigraphic scheme of the upper triassic – lower cretaceous trending sw–ne across the norwegian–danish basin. sequence stratigraphic key surfaces are included, numbered sequence boundaries are shown in red (sb) and maximum flooding surfaces shown in blue (mfs). the sbs bound the named sequences (e.g. vi 1, ga 1). the main tectonic events are indicated. note the significant base middle jurassic unconformity that cuts deep into older strata to the south-west and is onlapped as the area of subsidence expanded during late jurassic times. od, oddesund formation; rkf, ringkøbing–fyn high; sk, skagerrak formation; skp, skagerrak–kattegat platform; stz, sorgenfrei–tornquist zone. modified from nielsen (2003); time-scale from gradstein et al. (1994). 16 well, in which the succession reflects a transition to the contemporaneous skagerrak formation that dominates along the northern and eastern basin margin (bertelsen 1980; nielsen & japsen 1991). the bunter shale and bunter sandstone formations are present in felicia-1a, whereas the ørslev, falster and tønder formations are replaced by the skagerrak formation. the bunter sandstone formation may also be present in the terne-1 well where the lower 155 m of the triassic succession consists mainly of fine-grained, well-sorted sandstones. the bunter sandstone formation mainly consists of red-brown and yellow-brown, mediumto fine-grained, well-sorted sandstones with intraformational claystone clasts and thin mudstone beds, largely recording deposition in ephemeral, braided fluvial channels in an arid desert environment (bertelsen 1980; pedersen & andersen 1980). eolian dune sand, and mud deposited in ephemeral lakes, may constitute minor proportions of the formation. up-section and towards the northern and north-eastern basin margin, the bunter sandstone formation passes into the skagerrak formation. on the skagerrak–kattegat platform, lower –middle triassic strata are all referred to the skagerrak formation, which here seems to include large parts of the upper triassic as well (frederikshavn-1, -2, -3, sæby-1; nielsen & japsen 1991; figs 8, 9). the formation consists of a heterogenous succession of interbedded conglomerates, sandstones, siltstones and claystones that were mainly deposited as alluvial fans along the basin margins. late triassic clastic and evaporitic deposition and marine flooding in late triassic times, the arid or semi-arid climate continued and deposition of variegated red-brown or brown, calcareous, anhydritic and pyritic mudstones and siltstones with thin beds of dolomitic limestone and marl commenced in sabkhas and ephemeral lakes. in the central, deep parts of the basin, more permanent lakes were established. the deposits are included in the carnian – lower norian od desund formation, which passes into the skagerrak formation towards the basin margins to the north and north-east (figs 8, 9; bertelsen 1980). in places, the od desund formation includes two halite units up to 90 m thick, which in some areas have contributed to the formation of salt domes together with the zechstein salts (liboriussen et al. 1987; jensen & schmidt 1993; christensen & korstgård 1994). a gradual change to more humid conditions took place in late triassic times, associated with an early norian marine transgression that probably came from the south and resulted in the formation of a large epicontinental sea. the transgression led to deposition of oolitic limestones succeeded by marlstones and fossiliferous claystones of the vinding formation, which is typically 40–100 m thick over most of the basin (figs 8–10; bertelsen 1978, 1980; nielsen 2003). at its maximum extent in late norian times, the shallow sea covered most of the central basin and the ringkøbing–fyn high, whereas deposition of fluvial arkosic sands and lacustrine muds of the skagerrak formation continued in the sorgenfrei–tornquist zone and on the skagerrak–kattegat platform. maximum transgression was followed by phased regression, and shoreface and fluvial sands of the lower gassum formation were deposited in stepwise, more basinward positions; the sands are interbedded with clays of the upper vinding formation in the basin centre. deposition of fluvial sand and lacustrine mud of the skagerrak formation continued in the fjerritslev trough (figs 8, 10). regression culminated in the early rhaetian with the formation of an extensive, fluvially incised sequence boundary (sb 5 of nielsen 2003; figs 8, 10). the following, widespread marine flooding was initiated with deposition of fluvial–estuarine sediments, up to 30 m thick in the basin centre, above the sequence boundary. the transgression was punctuated by two shortterm, forced regressions that led to deposition of widespread shoreface sand sheets encased in offshore mud (hamberg & nielsen 2000; nielsen 2003). the transgression reached its maximum in the latest rhaetian, when the entire study area and the ringkøbing–fyn high were covered by the sea, and marine mudstones were deposited widely (mfs 7; figs 8, 10–11). hettangian – early pliensbachian transgression and basin expansion during latest rhaetian times, the climate changed to the subtropical to warm-temperate and humid conditions that characterised the jurassic period, when large quantities of clay were supplied to the basin due to weathering of pa laeozoic shales and granitic basement of the baltic shield. the jurassic transgression was interrupted by two phases of coastal progradation that caused deposition of two thin, regressive, shoreface sand sheets, which constitute the uppermost part of the gassum formation over much of the study area. the regression culminated in coastal progradation far into the basin accompanied by fluvial erosion and incision in the himmerland graben, the fjerritslev trough and the skagerrak–kattegat platform (sb 9; figs 10–11). sub sequently, the regional transgression continued (ts 9; early hettangian planorbis zone), so that fully marine mudstones belonging to the f-ia unit of the f-i member of the fjerritslev formation overlie the sandy gassum formation over most of the study area (figs 8–14). the mudstones have a high content of land-derived organic matter. the transgression peaked in the early and late hettangian (mfs 9, 10), interrupted by a short-term regression in the middle hettangian (sb 10; figs 8, 10–12). deposition of transgressive paralic deposits along the basin margin was interrupted briefly by a fall in sea level soon after the hettangian–sinemurian boundary. this resulted in fluvial 17 oddesund fm system series stage ju ra ss ic t ri as si c åsgard formation leek member bo member heno fm lola formation bryne formation norwegian– danish basin sw ne vedsted formation børglum formation flyvbjerg formation haldager sand formation fj er ri ts le v fo rm at io n gassum formation vinding formation skagerrak formation f-iv mb f-iii mb f-ii mb f-ib f-i mb f-ia frederikshavn formation danish central graben ryazanian volgian kimmeridgian oxfordian callovian bathonian bajocian aalenian toarcian pliensbachian sinemurian hettangian rhaetian norian winterton formation c re ta ce ou s valanginian fjerritslev formation vyl fm poul fmfarsund formation lo w er repp u m id dl e lo w er u pp er u l u l l u u m l l m u m l u l u u l l u m l u u l u m l paralic and non-marine sandstones, siltstones, mudstones and coals marine mudstones and siltstones unconformity offshore organic-rich marine shales submarine fan sandstones and siltstones shallow marine sandstones and siltstones, offshore mudstones alluvial sandstones and source rocks hiatus lacustrine mudstones marine calcareous claystones, carbonates middle graben formation lulu formation fig. 9. lithostratigraphy of the upper triassic – lower cretaceous in the norwegian–danish basin. the stratigraphy of the danish central graben is included for reference. modified from michelsen et al. (2003). 18 so rg en fr ei –t or nq ui st z on e 25 k m 37 k m 5 km 41 k m o dd es un d fo rm at io n o dd es un d fo rm at io n t s 7 h im m er la nd g ra be n ve ds te d1 sp m ej ru p1 g r r ød di ng -1 sp h yl le bj er g1 g r fa rs ø1 g r n e sw sb 2 ls t h st t st h st t st ls t h st t st fj 1 g a1 v i 1 v in di ng f m sk ag er ra k fm m fs 1 m fs 9 sb 9 t s 9 m fs 7 sb 8 sb 6 sb 4 m fs 8 sb 7 m fs 6 m fs 5 t s 5 sb 5 m fs 4 sb 3 m fs 3 t s 4 t s 1 t s 7 sb 6 sb 4 m fs 9 sb 9 t s 9 m fs 7 sb 8 m fs 8 sb 7 m fs 6 m fs 5 t s 5 sb 5 m fs 4 sb 3 m fs 3 t s 4 fl uv ia l es tu ar in e la cu st ri ne 50 m r in gk øb in g– fy n h ig h o ffs ho re sh or ef ac e m ej ru p1 r ød di ng -1 fa rs ø1 h yl le bj er g1 ve ds te d1 50 k m sy st em s tr ac ts ls t lo w st an d t st tr an sg re ss iv e h st h ig hs ta nd bo un di ng s ur fa ce s sb se qu en ce b ou nd ar y m fs m ax im um m ar in e flo od in g su rf ac e t s tr an sg re ss iv e su rf ac e v i 1 se qu en ce fj er ri ts le v fm g as su m f m d ep os iti on al e nv ir on m en ts ska ger rak –k att ega t pl atf orm fj er rit sle v tr ou gh h i. g r. n or w eg ia n– d an is h ba si n so rg en fre i– to rn qu ist z on e fi g. 1 0. w el l-l og p an el tr en di ng s w –n e a cr os s t he n or w eg ia n– d an ish b as in (s ee in se t m ap ; h i. g r. , h im m er la nd g ra be n) sh ow in g th e n or ia n– h et ta ng ia n de po sit io na l e nv iro nm en ts o f t he v in di ng , s ka ge rr ak , g as su m a nd lo w er m os t f je rr its le v fo rm at io ns . n ot e th at th e pr om in en t s eq ue nc e st ra tig ra ph ic su rf ac es (m fs 1, m fs 7, m fs 9, s b 5, s b 9) d ef in e m aj or se qu en ce s v i 1 , g a 1, a nd f j 1 ; t he se se qu en ce s a re su bd iv id ed in to m in or se qu en ce s b ou nd ed b y sb 3 et c. g r , g am m a ra y; s p, se lfpo te nt ia l. m od ifi ed fr om n ie lse n (2 00 3) . 19 m ej ru p1 g r h yl le bj er g1 g r ve ds te d1 sp fl yv bj er g1 sp fr ed er ik sh av n2 sp bø rg lu m -1 sp bø rg lu m fa ul t 10 0 m fj er ri ts le v tr ou gh h im m er la nd g ra be n sk ag er ra k– k at te ga t pl at fo rm 65 k m 41 k m 30 k m 15 k m 42 k m t s 7 t s 9 m fs 7 sb 1 2 sb 1 1 sb 8 t s 11 m fs 9 m fs 1 1 m fs 1 2 t s 12 sb 9 m fs 7 sb 1 1 sb 1 0 t s 10 sb 9 sb 8 t s 9 m fs 8 t s 11 t s 7 m fs 9 m fs 1 0 m fs 1 1 m fs 1 2 sb 1 2 sw n e ls t ls t t st ls t h st t st h st /f r st t st h st t st fr st ls t t st ls t ls t fr st h st t st t st h st ls t h st g a 1 fj 4 fj 3 fj 2 fj 1 fj er ri ts le v fm g as su m f m sk ag er ra k fm fj er ri ts le v fm d ep os iti on al e nv ir on m en ts fl uv ia l es tu ar in e la cu st ri ne la go on al sh or ef ac e o ffs ho re sy st em s tr ac ts ls t lo w st an d t st tr an sg re ss iv e h st h ig hs ta nd fr st fo rc ed r eg re ss iv e bo un di ng s ur fa ce s sb se qu en ce b ou nd ar y m fs m ax im um m ar in e flo od in g su rf ac e t s tr an sg re ss iv e su rf ac e fj 1 se qu en ce r in gk øb in g– fy n h ig h ve ds te d1 50 k m fl yv bj er g1 m ej ru p1 h yl le bj er g1 bø rg lu m -1 fr ed er ik sh av n2 ska ger rak – katt ega t platf orm fj er rit sle v tr ou gh h i. g r. n or w eg ia n– d an is h ba si n so rg en fre i– to rn qu ist z on e fi g. 1 1. w el l-l og p an el tr en di ng s w –n e ac ro ss th e n or w eg ia n– d an ish b as in (s ee in se t m ap ; h i. g r. , h im m er la nd g ra be n) sh ow in g th e r ha et ia n– si ne m ur ia n de po sit io na l e nv iro nm en ts o f t he u pp er g as su m a nd lo w er f je rr its le v fo rm at io ns . m od ifi ed fr om n ie lse n (2 00 3) . 20 m fs 1 5 sb 1 9– 22 m fs 2 2 sb 2 3 m fs 1 2 sb 1 5 sb 1 3 sb 1 2 m fs 1 1 sb 1 1 sb 9 m fs 7 m fs 1 sb 5 m fs 9 sk ag en -2 sæ by -1 fr ed er ik sh av n2 bø rg lu m -1 fl yv bj er g1 ve ds te d1 sw n e h im m er la nd g ra be n so rg en fr ei –t or nq ui st z on e sk ag er ra k– k at te ga t pl at fo rm m fs 2 2 sb 1 6 m fs 1 7 sb 1 7 m fs 1 6 h a 3fl 1 35 k m 10 k m 34 k m 15 k m 30 k m sp sp sp sp sp g r sb 1 7 fj 9 10 0 m h a 1 h a 2 sb 5 m fs 7 m fs 9 sb 9 sb 1 0 sb 1 1 sb 1 2 m fs 1 2 m fs 1 1 m fs 1 0 sb 1 3 m fs 1 3 m fs 1 4 sb 1 5 sb 1 4 sb 1 6 sb 1 9 sb 2 0 sb 2 1 sb 2 2 sb 2 3 t s 19 m fs 1 6 m fs 1 5 fj 8 g a 1 fj 1 fj 2 fj 3 fj 6 fj 5 fj 4 fj 7 v i 1 bø rg lu m fa ul t 50 k m fl yv bj er g1 sk ag en -2 bø rg lu m -1 fr ed er ik sh av n2 sæ by -1 fj er ri ts le v fm bø rg lu m f m h al da ge r sa nd f m fl yv bj er g fm g as su m fm sk ag er ra k fm sk ag er ra k– katt eg at pla tfo rm fj er rit sle v tr ou gh n or w eg ia n– d an is h ba si n so rg en fre i– to rn qu ist z on e ve ds te d1 fl uv ia l es tu ar in e la cu st ri ne la go on al sh or ef ac e o ffs ho re d ep os iti on al e nv ir on m en ts bo un di ng s ur fa ce s sb se qu en ce b ou nd ar y m fs m ax im um m ar in e flo od in g su rf ac e t s tr an sg re ss iv e su rf ac e r in gk øb in g– fy n h ig h h i. g r. g a 1 s eq ue nc e fi g. 1 2. w el l-l og p an el tr en di ng s w –n e a cr os s t he n or w eg ia n– d an ish b as in (s ee in se t m ap ; h i. g r. , h im m er la nd g ra be n) sh ow in g th e n or ia n– o xf or di an d ep os iti on al e nv iro nm en ts o f t he s ka ge rr ak , g as su m , f je rr its le v, h al da ge r s an d, f ly vb je rg a nd b ør gl um f or m at io ns . n ot e th e pr on ou nc ed e ro sio na l t ru nc at io n at th e am al ga m at ed se qu en ce b ou nd ar y, s b 1 9– 22 . m od ifi ed fr om n ie lse n (2 00 3) . 21 10 0 m m ej ru p1 h yl le bj er g1 ve ds te d1 bø rg lu m -1 fl yv bj er g1 fr ed er ik sh av n2 g r g r so ni c so ni c sp r es sp r es sp r es sp r es flv flv flll fll flb flb flll fll flb fla fla fl yv b. –h al d. fj er ri ts le v fm bø rg lu m f m fjerritslev fm m ar in e m ud st on es sh al lo w m ar in e sa nd st on es a nd s ilt st on es pa ra lic a nd n on -m ar in e sa nd st on es , si lts to ne s, m ud st on es a nd c oa ls n or w eg ia n– d an is h ba si n so rg en fr ei –t or nq ui st z on e sk ag er ra k– k at te ga t pl at fo rm sk ag er ra k fm (t ri as si c) fr ed er ik sh av n fm bø rg lu m f m ve ds te d fm g as su m f m (t ri as si c) flll fll flb fla g as su m f m (ju ra ss ic ) fr ed er ik sh av n fm fl yv bj er g fm h al da ge r sa nd f m ve ds te d fm flv flll fll flb ve ds te d1 50 k m fl yv bj er g1 m ej ru p1 r ød di ng -1 h yl le bj er g1 bø rg lu m -1 sk ag er rak – katt eg at pla tfo rm fr ed er ik sh av n2 h i. g r. fj er ri ts le v tr ou gh n or w eg ia n– d an is h ba si n so rg en fre i– to rn qu ist z on e ri ng kø bi ng –f yn h ig h fi g. 1 3. w el l-l og p an el tr en di ng s w –n e a cr os s t he b as in (s ee in se t m ap ; h i. g r. , h im m er la nd g ra be n) sh ow in g th e lit ho st ra tig ra ph y of th e u pp er t ri as sic – l ow er c re ta ce ou s. m od ifi ed fr om m ic he lse n et a l. (2 00 3) . 22 incision on the skagerrak–kattegat platform, while regressive shoreface sand was deposited in the fjerritslev trough (sb 11; figs 8, 11–12). farther basinwards, heteroliths and silty mudstones were deposited above the conformable part of the sequence boundary. a rapid sea-level rise followed in the earliest sinemurian (upper part of the bucklandi zone), and transgressive marine muds of the f-ib unit of the f-i member of the fjerritslev formation finally overstepped fluvial and marine sands of the gassum formation in the sorgenfrei–tornquist zone and on the skagerrak–kattegat platform (figs 8, 9, 13, 14). the early sinemurian transgression led to deposition of up to 150 m of uniform mudstones in the basin, showing a marked thinning towards the northern and north-eastern basin margin (sequence fj 3 in figs 8, 11, 12). following a minor sea-level fall in the late sinemurian, the overall early jurassic sea-level rise continued and reached a maximum in the latest sinemurian (fig. 14). in the centre of the basin, the diversity and abundance of the ostracod fauna decreased and infaunal bivalves and some of the epifaunal bivalves disappeared due to reduced oxygen conditions (pedersen 1986; michelsen 1989a). a gradual decrease in the rate of sea-level rise in the early pliensbachian jamesoni zone caused a distinct ba sinward progradation of shoreface clinoform sandstones on the skagerrak–kattegat platform. the regression culminated in the middle early pliensbachian (early ibex zone) and deposition changed from fine-grained mud (f-ib unit, f-i member) to silty and sandy heteroliths (f-iia unit, f-ii member, sb 13; figs 8, 9, 12–14; michelsen 1989a; nielsen 2003). when the sea level started to rise again, deposition of fine-grained mud resumed in the ba sin (lower part of f-iib unit, f-ii member, fjerritslev formation), while backstepping parasequences of marine sand were succeeded by transgressive mud on the ska gerrak–kattegat platform. peak transgression was reached in the late early pliensbachian davoei zone. thereafter, the rate of sea-level rise decreased and a coarsening-upward succession of mud and fine-grained heteroliths was deposited in the study area (middle part of f-iib unit, f-ii member). late pliensbachian – early aalenian sea-level fluctuations significant erosion took place on the skagerrak–kattegat platform during a sea-level fall in the early late pliensbachian margaritatus zone (sb 14; figs 8, 12). basinwards, deposition changed to silty and sandy mud and fine-grained sand, showing pronounced thinning over some salt structures possibly reflecting shallow-water depths (upper part of f-iib and f-iic beds, f-ii member; michelsen 1989a, b). the ensuing sea-level rise reached a peak in the late late pliens bachian (early spinatum zone). marine silty mud accumulated in the basin, while muddy marine sand with bivalves was deposited in the fjerritslev trough and presumably also in the farsund basin (lower part of the f-iii member, fjerritslev formation; figs 8, 9, 13). strata from this period are absent on the skagerrak–kattegat platform due to bypass or later erosion (fig. 8). a subsequent sea-level fall caused the formation of a widespread marine regressive surface of erosion and deposition of 5–10 m of lowstand shoreface sandstones in the central parts of the fjerritslev trough (sb 15; figs 8, 12); lower pliensbachian strata were eroded on the ska gerrak–kattegat platform. the subsequent sea-level rise caused marine flooding over the entire study area, and the transgression reached its maximum in the early toarcian falciferum zone (mfs 15; figs 8, 12, 14). due to oxygen-poor conditions, the ostracod fauna disappeared and an increasing amount of algalderived marine organic matter was preserved, resulting in the accumulation of organic-rich, oil-prone mudstones in parts of the basin (f-iii and f-iv members, see below; figs 8, 9, 13). during the remainder of the early jurassic and in the early aalenian opalinum zone (early middle jurassic), a succession of up to 150 m of marine mudstones was deposited in the sorgenfrei–tornquist zone (f-iii and f-iv members, fjerritslev formation; figs 8, 9, 12–14). interbedded with the mudstones are three shoreface sandstones, 5–15 m thick, overlying regressive marine erosion surfaces. these sandstones were deposited during sea-level falls and could act as carrier beds for any petroleum generated and expelled from thermally mature toarcian source rocks. late early – middle jurassic uplift and erosion the ringkøbing–fyn high and most of the nor we gian–danish basin were uplifted in late early jurassic – early middle jurassic times, and the triassic – lower jurassic successions were eroded on the highest parts of the ringkøbing–fyn high (figs 8, 14, 15). the lower jurassic succession, including the potential source-rock intervals of the fjerritslev formation, was deeply eroded in the up lifted area north of the high. outside the study area, to wards the north-west, norwegian wells (10/7-1, 10/8-1, 11/9-1, 11/10-1; fig. 1) confirm the deep regional erosion of the lower jurassic succession, although the large hiatuses recorded in these wells may not be representative 23 since they intersect positive jurassic structures that were exposed to deep erosion during regional uplift. in the egersund basin, located on-strike farther to the north-west (fig. 1), the fjerritslev formation is also deeply truncated and is locally absent. erosion did not reach such deep levels closer to the sorgenfrei–tornquist zone (fig. 15). within the fault-bounded sorgenfrei–tornquist zone itself, where subsidence still occurred (but at a much lower rate than before), the dramatic changes in regional basin configuration are marked by a shift from offshore mudstones 4°e 8°e 12°e 16°e 58°n 56°n late aalenian – bathonian 4°e 8°e 12°e 16°e 58°n 56°n hettangian 4°e 8°e 12°e 16°e 58°n 56°n sinemurian – early pliensbachian 4°e 8°e 12°e 16°e 58°n 56°n toarcian – early aalenian sweden norway germanythe netherlands sweden norway germanythe netherlands sweden norway germanythe netherlands sweden norway germany denmark denmark the netherlands 100 km 100 km 100 km 100 km denmark denmark a b c d denmark offshore marine (mud-dominated) deep marine (sands, gravels) non-deposition/erosion paralic and non-marine sandstones, siltstones, mudstones and coals shallow marine sandstones and siltstones 4°e 8°e 12°e 16°e 58°n 56°n callovian 4°e 8°e 12°e 16°e 58°n 56°n kimmeridgian 4°e 8°e 12°e 16°e 58°n 56°n volgian – early ryazanian 100 km sweden norway germany the netherlands sweden norway germany the netherlands sweden norway germany the netherlands 100 km 100 km denmark denmark e g f fig. 14. paleogeographic maps showing changes in the general depositional environments through jurassic times. modified from michelsen et al. (2003). 24 of the f-iv member (fjerritslev formation) to shallow marine sandstones (haldager sand formation; figs 8, 12–16). during the rest of the aalenian, the bajocian and the early bathonian, deposition was more or less confined to the narrow zone bounded by the fjerritslev and børglum faults and their south-eastward continuation in the kattegat, øresund and southern sweden (figs 8, 14). it is assumed that the westward continuation of the faults defining the southern margin (the fjerritslev fault) and the northern margin of the farsund basin also delineate an area of continued subsidence in middle jurassic times, thus protecting the potential source rocks in the upper fjerritslev formation from erosion in that area. the fjerritslev trough and presumably also the farsund basin received detritus from the uplifted areas to the west and south-west, and from the baltic shield to the north, resulting in deposition of the haldager sand formation, which is 29–155 m thick in the well sections in the fjerritslev trough (figs 8, 9, 16). the formation is dominated by shoreface and fluvial sandstones interbedded with thin marine and lacustrine mudstones, and thin coaly beds in places. late middle – late jurassic basin expansion regional subsidence resumed during late middle – late jurassic times and the shallow marine clastic depocentre gradually expanded. the upper jurassic thus onlaps onto the base middle jurassic unconformity, showing significant younging of the onlap towards the south-west and northeast, on both sides of the sorgenfrei–tornquist zone (figs 14, 17). deposition overstepped the bounding faults of the sorgenfrei–tornquist zone, with deposition of bath onian(?) braided fluvial sands on the skagerrak–kattegat platform and in the himmerland graben. deposition also resumed outside the study area where bathonian(?) sandstones of the haldager sand and bryne formations are preserved in the 10/7-1 and 10/8-1 wells. a marine transgression close to the callovian–oxfordian boundary influenced most of the basin, and accommodation space was also created in the former by-pass zone of the southern part of the basin and on the skagerrak–kattegat platform, where fluvial sands were deposited. during the u. toarcian l. toarcian ? ? ? l. sin em urian u. si nem uri an norway sweden ? hettangian hettan gian 6°e 10°e 58°n 56°n subcrop to the base cretaceous unconformity fault distribution of f-iii/f-iv mbs of the fjerritslev fm well intra-aalenian basinward shift in facies middle–upper jurassic absent subcrop to the base middle jurassic unconformity 100 km the netherlands germany u. pliensbach. l. pliensbach. sinemurian hettangian u. triassic fig. 15. the subcrop below the base middle jurassic unconformity and the distribution of the stratigraphic interval with the lower jurassic potential source rocks (f-iii and f-iv members). the coarse-dotted area indicates the area within which the shift from marine mudstones to shallow marine or fluvial sandstones occurred in the aalenian without major erosion. modified from nielsen (2003). 25 fl yv bj er g fm t s 23 m fs 2 2 sb 2 3 t s 23 t s 21 t s 20 t s 19 m fs 2 2 m fs 2 0 m fs 1 7 m fs 1 6 sb 2 3 sb 2 2 t s 22 sb 1 6 m fs 1 6 14 5 km 14 8 km 32 k m 35 k m te rn e1 g r h al da ge r1 sp fl yv bj er g1 sp fr ed er ik sh av n2 sp sk ag en -2 sp bø rg lu m f m fj er ri ts le v fm h al da ge r sa nd f m t s 22 so rg en fr ei –t or nq ui st z on e sk ag er ra k– k at te ga t pl at fo rm sw n e fl 1 bø 1 h a 3 h a 2 h a 1 fj 1 0 fj 9 fj 8 fj 7 t st ls t ls t ls t t st ls t t st ls t t st h st h st h st t st t st ls t ls t h st t st t st h st sb 2 2 sb 2 1 sb 2 0 sb 1 9 sb 1 8 sb 1 7 sb 1 6 bø rg lu m fa ul t 10 0 m 50 k m fl yv bj er g1 fr ed er ik sh av n2 sk ag en -2 te rn e1 h al da ge r1 sk ag er rak – katt eg at pla tfo rm fj er rit sle v tr ou gh h i. g r n or w eg ia n– d an ish b as in so rg en fre i– to rn qu ist z on e m fs 1 8 fl uv ia l es tu ar in e la go on al sh or ef ac e o ffs ho re d ep os iti on al e nv ir on m en ts bo un di ng s ur fa ce s sb se qu en ce b ou nd ar y m fs m ax im um m ar in e flo od in g su rf ac e t s tr an sg re ss iv e su rf ac e h a 1 s eq ue nc e m fs 1 9 sy st em s tr ac ts ls t lo w st an d t st tr an sg re ss iv e h st h ig hs ta nd fi g. 1 6. w el l-l og p an el tr en di ng s w –n e a cr os s t he n or w eg ia n– d an ish b as in (s ee in se t m ap ; h i. g r. , h im m er la nd g ra be n) sh ow in g th e to ar ci an –o xf or di an d ep os iti on al e nv iro nm en ts o f t he u pp er f je rr its le v, h al da ge r s an d, f ly vb je rg an d b ør gl um f or m at io ns . n ot e th e va ry in g se ve ri ty o f e ro sio na l t ru nc at io n at th e ba se o f t he m id dl e ju ra ss ic d ep en di ng on st ru ct ur al p os iti on , a s r ef le ct ed b y th e tr un ca tio n of th e se qu en ce s d ef in ed in th e lo w er ju ra ss ic su cc es sio n. m od ifi ed fr om n ie lse n (2 00 3) . 26 oxfordian, the sedimentation area was further enlarged, and a north-eastwards thickening wedge of transgressive, fossiliferous marine sand and mud was deposited above lagoonal deposits on the skagerrak–kattegat platform (lower flyv bjerg formation; figs 8, 9, 13, 16). the transgression peaked in the mid-oxfordian, with deposition of marine mudstones over most of the northern part of the study area including the fjerritslev trough and the skagerrak–kattegat platform. a sea-level fall in the latest oxfordian resulted in coastal progradation on the skagerrak–kattegat platform and in the fjerritslev trough; fluvial and shallow marine sands were deposited, and a south-west prograding wedge was formed (upper flyvbjerg formation; figs 8, 9, 16). extensive marine flooding occurred in kimmeridgian times, and sedimentation of marine mud (børglum for mation) characterised the whole area, although marked thinning towards the south-west of the study area emphasises the reduced accommodation space there (figs 8, 12–14, 16). during volgian–ryazanian times, the depositional environment was dominantly a shallow shelf, with three to four major phases of coastal progradation (sequences fr 1–3 of the frederikshavn formation; figs 8, 14). marine muds were deposited over much of the study area. cretaceous continued basin expansion and chalk deposition in early cretaceous times, the area of marine deposition expanded further with coastal progradation from the north and north-east. deposition of marine mud prevailed over most of the study area. the lower cretaceous mudstones with sandy intercalations (most common towards the northeast) are included in the vedsted formation (fig. 9), which consists of four depositional units (michelsen & nielsen 1991). in late cretaceous – danian times, a high sea level dominated and the study area was covered by an epicontinental sea. the dry climate and low relief of the hinterland reduced clastic input, and biogenic, pelagic chalk deposition (with coccolith plates being the dominant constituent) occurred over the entire study area. in the easternmost part of the basin and within the fennoscandian border zone, deposition of marine greensands ceased in late cenomanian times and in the early turonian intermittent deposition of marls and mudstones in parts of the basin also came to an end (stenestad 1972; surlyk 1980). in central parts of the study area, in the deepest parts of the cretaceous basin, coc6°e 10°e 14°e 58°n 56°n ? ? ? ? bajocian–bathonian oxfordian oxfordian sweden norway the netherlands germany onlap limits uncertain fault well intra-aalenian basinward shift in facies middle–upper jurassic absent onlap limits 100 km kimmeridgian–volgian fig. 17. map showing the upper jurassic – lower cretaceous onlap onto the base middle jurassic unconformity, indicating the gradual expansion of the depositional area. modified from nielsen (2003). 27 colith-dominated chalks accumulated in water depths that may have been up to 500–600 m (surlyk & lykke-andersen 2007). closer to the basin margins and over structural highs, chalks rich in bryozoans and other benthic fossils were deposited at mid to inner shelf depths (?100–200 m). in more shallow water areas in the fennoscandian border zone, benthos-rich chalks pass into bryozoan wackestones and packstones that locally developed as mound complexes, while skeletal grainstones and oyster bank carbonates formed closer to the shoreline (surlyk 1997). centrally in the study area, south-west of the sorgen frei–tornquist zone, 1.5–2 km of chalk was deposited, while 500–750 m accumulated over the ringkøbing–fyn high. in the sorgenfrei–tornquist zone, the original thickness of the chalk succession is masked by late cretaceous – palaeogene inversion and erosion (liboriussen et al. 1987; nielsen & japsen 1991; jensen & schmidt 1993; michelsen & nielsen 1993; erlström & sivhed 2001). the inversion began in coniacian times and accelerated rapidly during santonian–campanian times (the sub-hercynian phase; ziegler 1990); quiescence during maastrichtian–danian times was followed by pronounced inversion again in the late paleocene (the laramide phase; liboriussen et al. 1987; ziegler 1990). cenozoic clastic deposition and neogene exhumation after cessation of carbonate deposition in the paleocene, deep marine sedimentation of fine-grained hemipelagic deposits took over in the study area. the northern and eastern limits of these fine-grained sediments are unknown due to later erosion. in the oligocene, major clastic wedges began to build out from the baltic shield, while prodeltaic glauconite-rich clayey sediments were deposited farther basinwards. coarse-grained sediments reached the southern part of the basin and the ringkøbing–fyn high in neogene times (larsen & dinesen 1959; friis et al. 1998; rasmussen 2004). deposition in the greater north sea basin continued during the pliocene, and up to 500 m of sediments were deposited in the norwegian–danish basin during the late miocene and pliocene (overeem et al. 2001). during the post-late cretaceous period, parts of the norwegian–danish basin and the fennoscandian border zone were uplifted and eroded. this major tilting continued into the quaternary (japsen 1993; jensen & schmidt 1993; japsen et al. 2002a, b; rasmussen et al. 2005). 28 due to late cretaceous – early cenozoic inversion of fault blocks in the sorgenfrei–tornquist zone and neo gene–pleistocene regional uplift of the norwegian–danish basin (e.g. michelsen & nielsen 1991, 1993; jensen & schmidt 1993; japsen 1993, 1998; petersen et al. 2003a), present-day burial depths must be corrected for post-early cretaceous net exhumation (fig. 18). petersen et al. (2003a) presented a regional coalification curve for the norwe gian–danish basin based on 249 measurements from 15 wells (onshore wells and the hans-1 well in the kattegat). that study did not include offshore wells from the skagerrak area. the well-sections were corrected for net-exhumation values obtained from the analysis of sonic velocities of shales by japsen (1993). the accuracy of this method is dependent on a uniform shale unit covering the entire study area and a valid sonic velocity reference curve. net exhumations for wells not included in japsen’s (1993) study were estimated by comparison to nearby wells and interpolation. in the present study, new regional coalification curves for the norwegian–danish basin have been constructed in order to evaluate the depth to the oil window. the depths of the samples have been corrected for net exhumation, the magnitudes of which have been derived from both shale and chalk velocities. a total of 560 vitrinite reflectance (vr) measurements from 26 wells in the norwegian–danish basin were available for construction of the coalification curves (fig. 18). the vr values are from thomsen (1980, 1983), schmidt (1985, 1988, 1989) and geus unpublished data. all vr values are random measurements performed on core samples, sidewall cores or cuttings, and as many particles as possible were measured in each sample. untreated rock samples (whole rock) were used, as such samples – compared to kerogen concentrates – have the advantage that it is easier to identify the primary vitrinite particles and thus avoid oxidised and bituminous organic matter (e.g. barker 1996). identification of the primary, i.e. indigenous or autochthonous, vitrinite is essential for obtaining reliable vr values as this vitrinite reflects the actual therregional coalification curves corrected for net exhumation table 1. net exhumation magnitudes* well shale† chalk† mean§ vr comments anholt-1 1400 1400 comparison to hans-1, terne-1 and frederikshavn-1 (very uncertain) års-1 442 461 børglum-1 1185 600 c-1 300 306 comparison to inez-1 and vemb-1 (264 m) d-1 50 50 comparison to l-1 f-1 481 357 419 1200 farsø-1 377 530 felicia-1/1a 1020 712 866 800 fjerritslev-2 1443 802 frederikshavn-1 1000 702 comparison to sæby-1 (1051 m) gassum-1 1090 579 comparison to voldum-1 (845 m) + 250 m deeper chalk truncation haldager-1 1400 486 comparison to børglum-1 and fjerritslev-2 hans-1 1735 hobro-1 450 543 comparison to års-1 and kvols-1 hyllebjerg-1 470 552 inez-1 344 445 395 850 k-1 624 414 519 1300 kvols-1 361 420 l-1 0 0 mors-1 690 602 r-1 150 176 comparison to c-1, l-1, s-1 (0 m) and vemb-1 (238 m) rønde-1 394 447 skagen-2 1000 1000 comparison to børglum-1, frederikshavn-1 and sæby-1 (1051 m) terne-1 1373 vedsted-1 1300 500 comparison to børglum-1, fjerritslev-2 and haldager-1 vinding-1 175 250 comparison to mejrup-1 (sh: 231 m; ch: 253 m) and vemb-1 (sh: 264 m; ch: 238 m) * net exhumation magnitudes based on shale and chalk sonic velocity data (japsen et al. 2007) † magnitudes in italics are estimated; see comments column § mean of shale (sh) and chalk (ch) 29 mal maturity of the organic matter at the sampled depth. higher reflecting vitrinite particles may represent recycled organic matter that attained higher maturity from the temperature history of its former host rock (e.g. bostick 1979; hunt 1996; taylor et al. 1998). in contrast, vitrinite that yields anomalously low reflectance values may be suppressed (buiskool toxopeus 1983; carr 2000a, b). a more detailed description of the prerequisites for construction of a regional coalification curve is presented in petersen et al. (2003a). shale velocity-based curve the shale sonic velocity reference curve of japsen (1993) has recently been refined and new, modified net-exhumation values based on shale velocities have been proposed (table 1; japsen et al. 2007). the revised magnitudes of net exhumation are reduced by about 100 m compared to the values used by petersen et al. (2003a). the new net-exhumation values have been used to revise the coalification curve 0.4 0.60.2 0.8 1.0 1.2 1.4 1.6 anholt-1 års-1 børglum-1 c-1 d-1 f-1 farsø-1 felicia-1/1a fjerritslev-2 frederikshavn-1 gassum-1 haldager-1 hans-1 hobro-1 hyllebjerg-1 inez-1 k-1 kvols-1 l-1 mors-1 r-1 rønde-1 skagen-2 terne-1 vedsted-1 vinding-1 6000 4000 2000 0 d ep th pr es en t da y (m ) vitrinite reflectance (%ro) n = 560 a fig. 18. a: vitrinite reflectance values plotted against present-day depths for 26 onshore and offshore wells. 30 for the study area. in fig. 19, the present depths of samples from 25 wells have been corrected for the magnitudes of net exhumation obtained from shale velocities. for some wells, net-exhumation values based on shale velocities are not available, and the amount of net exhumation for these wells has been estimated by comparison to nearby wells and stratigraphic evaluation (table 1). despite the vr data showing considerable scatter, 13 of the wells define a wellconstrained vr trend (fig. 20a). following petersen et al. (2003a), the fjerritslev-2, haldager-1 and vedsted-1 wells from the fjerritslev trough are not included as they define an atypically steep maturation gradient. this was explained by a probable overestimation of shale velocities in this area due to an increased coarse-grained component within the lowermost fjerritslev formation. the anholt-1 and ter ne-1 wells show unusually low vr values and are thus not included. compared to the curve in petersen et al. (2003a), the horsens-1, lavø-1 and ullerslev-1 wells have been 0.2 0.3 0.5 0.70.4 0.6 0.8 4000 3000 2000 1000 0 vitrinite reflectance (%ro)b d ep th pr es en t da y (m ) n = 545 fig. 18. b: as in 18a, but excluding the felicia-1/1a deep well; for legend, see fig. 18a. 31 omitted, as only one vr measurement is available from each well. the vr data from the børglum-1 well are, however, included in the new coalification profile. accurately determined vr values will define a straight line in a semi-log plot (dow 1977), and the established vr profile yields a correlation coefficient of 0.89 (fig. 20a). the linear regression line intercepts the surface at a vr of c. 0.26%ro which is at the upper end of the reflectance values recorded for peaty organic matter at the surface (c. 0.10–0.25%ro; cohen et al. 1987). if the start of the oil window is set at a vr of c. 0.6%ro, the ‘shale curve’ suggests that the top of the oil window occurs at a burial depth of c. 3050 m. vr data from offshore wells in the skagerrak are not included in the coalification curve (fig. 20a). generally, the vr values from these wells lie above the maturity curve (fig. 19), and data from four of the wells (d-1, f-1, inez-1 and k-1) provide a relatively well-defined vr trend (fig. 20b). the regression line yields a correlation coefficient of 0.80 and intercepts the surface at a vr of c. 0.28%ro. this alternative maturity gradient may be applied for the off0.2 0.4 0.6 0.80.3 0.5 0.7 4000 3000 2000 1000 0 års-1 anholt-1 børglum-1 c-1 d-1 f-1 farsø-1 fjerritslev-2 frederikshavn-1 gassum-1 haldager-1 hans-1 hobro-1 hyllebjerg-1 inez-1 k-1 kvols-1 l-1 mors-1 r-1 rønde-1 skagen-2 terne-1 vedsted-1 vinding-1 vitrinite reflectance (%ro) d ep th co rr ec te d (s ha le ) ( m ) n = 545 fig. 19. vitrinite reflectance values from 25 wells plotted against depths corrected for post-early cretaceous uplift on the basis of shale sonic velo city data. 32 4000 3000 2000 1000 0 års-1 børglum-1 farsø-1 frederikshavn-1 gassum-1 hans-1 hobro-1 hyllebjerg-1 kvols-1 mors-1 rønde-1 skagen-2 vinding-1 10.90.80.70.60.50.40.30.260.2 vitrinite reflectance (%ro) d ep th co rr ec te d (s ha le ) ( m ) r2 = 0.89 n = 263 a 3000 2000 1000 0 r2 = 0.80 n = 166 d-1 f-1 inez-1 k-1 10.90.80.70.60.50.40.3 0.28 0.2 b d ep th co rr ec te d (s ha le ) ( m ) vitrinite reflectance (%ro) fig. 20. a: regional coalification curve for the norwegian–danish basin based on 12 onshore wells and the hans-1 well, selected from the wells in fig. 19 (see text). a total of 263 vitrinite reflectance values have been used. the regres sion line has a correlation coefficient of r2 = 0.89 and the line intercepts the surface at 0.26%ro. the depth to the top of the oil window at 0.6%ro is c. 3050 m. b: coalification curve based on four wells from the skagerrak area. the regression line yields a corre lation coefficient of 0.80 and the line intercepts the surface at a vr of c. 0.28%ro. the top of the oil window (vr c. 0.6%ro) occurs at a burial depth of c. 2600 m according to this curve. 33 shore part of the study area. according to this curve, however, the top of the oil window (vr c. 0.6%ro) occurs at a burial depth of c. 2600 m which, given the absence of petroleum discoveries in this part of the basin, is probably too shallow. underestimation of the magnitude of exhumation in the area may provide an explanation for the apparently erroneous curve. chalk velocity-based curve chalk velocities have – like shale velocities – been used to estimate the magnitude of net exhumation by japsen (1998). considerable uncertainty in the estimated exhumation may be introduced if the chalk section is <300 m thick. netexhumation magnitudes based on new chalk velocities (japsen et al. 2007) have been used to correct the presentday sample depths in this study (table 1). compared to the net-exhumation magnitudes in japsen (1998), the adjusted values presented here are reduced by about 200 m. in fig. 21, the present depths of samples from 25 wells have been corrected for the magnitudes of net exhumation obtained from chalk velocities. the vr data show a relatively large scatter, although the data seem to group into two elongate populations with the upper population principally formed by offshore wells in the skagerrak. the linear regression line for a regional vr curve based on 15 onshore wells and the hans-1 well has a correlation coefficient of 0.87 (fig. 22a), which is slightly poorer than the ‘shale curve’. the linear regression line is, however, based on more data (n = 282) as vr data from the fjerritslev-2 and vedsted-1 wells are included in the coalification profile. like the ‘shale curve’, the haldager-1 well yields an abnormally steep maturity gradient, while the anholt-1 and terne-1 wells display much too low vr values compared to the overall trend defined by the majority of wells (fig. 21). the curve intercepts the surface at a vr of c. 0.28%ro, which is at the upper end of the vr values of peaty organic matter (fig. 22a). compared to the ‘shale curve’, the slightly lower gradient of the ‘chalk curve’ yields a burial depth of c. 3100 m for the top of the oil window using a vr of 0.6%ro. as with the shale velocity-corrected samples, the chalk velocity-corrected samples from offshore wells in the skagerrak yield vr values that generally lie above the established maturity profile (fig. 21). three of these wells (f-1, inez-1, k-1) provide a well-constrained vr gradient (correlation coefficient of 0.90) that intercepts the surface at c. 0.26%ro (fig. 22b). according to this curve, the top of the oil window is located at about 2450 m depth, which is considered an unrealistically shallow burial depth probably due to underestimation of the magnitude of exhumation. depth to the top of the oil window if the top of the oil window is set at a maturity level corresponding to c. 0.6%ro, the two depth-corrected maturity gradients, derived from the onshore wells, yield similar depths to the top of the oil window (c. 3050–3100 m). according to the relationship between vr and burial peak temperature of barker & pawlewicz (1994), c. 0.6%ro corresponds to c. 95°c. this fits reasonably well with the pres ent-day thermal gradient of 31.4°c/km in felicia-1/1a, as estimated from corrected bottom-hole temperature measurements. in contrast, depth-corrected maturity gradients based on offshore wells from the skagerrak yield unrealistically shallow burial depths to the top of the oil window. burial depths of only 2450–2600 m would imply the presence of mature source rocks in the area, which is inconsistent with the lack of direct evidence of generated petroleum. 34 0.2 0.3 0.5 0.70.4 0.6 0.8 4000 3000 2000 1000 0 f ig. 21 års-1 anholt-1 børglum-1 c-1 d-1 f-1 farsø-1 fjerritslev-2 frederikshavn-1 gassum-1 haldager-1 hans-1 hobro-1 hyllebjerg-1 inez-1 k-1 kvols-1 l-1 mors-1 r-1 rønde-1 skagen-2 terne-1 vedsted-1 vinding-1 vitrinite reflectance (%ro) d ep th co rr ec te d (c ha lk ) ( m ) n = 545 fig. 21. vitrinite reflectance values from 25 wells plotted against depths corrected for post-early cretaceous uplift on the basis of chalk sonic velo city data. 35 års-1 børglum-1 farsø-1 fjerritslev-2 frederikshavn-1 gassum-1 hans-1 hobro-1 hyllebjerg-1 kvols-1 mors-1 rønde-1 skagen-2 vedsted-1 vinding-1 r2 = 0.87 n = 282 4000 3000 2000 1000 0 10.90.80.70.60.50.40.30.280.2 vitrinite reflectance (%ro) d ep th co rr ec te d (c ha lk ) ( m ) a f-1 inez-1 k-1 r2 = 0.90 n = 157 3000 2000 1000 0 10.90.80.70.60.50.40.3 0.26 0.2 b d ep th co rr ec te d (c ha lk ) ( m ) vitrinite reflectance (%ro) fig. 22. a: regional coalification curve for the norwegian–danish basin based on 14 onshore wells and the hans-1 well, selected from the wells in fig. 21 (see text). a total of 282 vitrinite reflectance values have been used. the regression line has a correlation coefficient of r2 = 0.87 and the line intercepts the surface at 0.28%ro. the depth to the top of the oil window at 0.6%ro is c. 3100 m. b: coalification curve based on three offshore wells from the skagerrak area. the vr data form a very wellconstrained vr gradient (correlation coefficient of 0.90) that intercepts the surface at c. 0.26%ro. according to this curve, the top of the oil window is located at only c. 2450 m depth. 36 the regional petroleum generation potential and thermal maturity of the pre-upper cretaceous succession in the study area have been assessed by the evaluation of more than 4000 data points (s1 and s2 yields, hydrogen index (hi), tmax, total organic carbon (toc)). these are derived from rock-eval pyrolysis and toc determination of principally cuttings samples from 33 wells situated onshore denmark and offshore in the skagerrak and kattegat areas (fig. 1). most of the wells drilled the lower cretaceous – triassic succession, but pre-permian and permian strata were also encountered in some wells. unreliable tmax or hi values derived from low s2 yields (i.e. approximately <0.2 mg hc/g rock) or toc contents have been omitted from the data set. the following thresholds for hi have been applied to evaluate the petroleum generation capacity of the predominantly type ii kerogen dominated source rocks (peters & cassa 1994): hi <200 mg hc/g toc: gas-prone hi = 200–300 mg hc/g toc: mixed oil/gas-prone hi >300 mg hc/g toc: oil-prone determination of the thermal maturity is based on vitrinite reflectance (vr) and tmax values, and the potential source rock is considered immature for vr values <0.6%ro or tmax values <435°c. the tmax range 435–460°c defines the oil window (bordenave et al. 1993; peters & cassa 1994). tmax is, however, also influenced by kerogen type and the mineral matrix, and single tmax values may therefore be less reliable as a maturity parameter (peters 1986). average tmax values for specific formations or members with source rocks are therefore used. pre-permian units seven wells have encountered pre-permian strata in the danish part of the norwegian–danish basin and fennoscandian border zone. source-rock data are available from five of them: frederikshavn-1, nøvling-1, rønde-1, slagelse-1 and terne-1 (fig. 1). bitumen and oil stains have been reported from lower palaeozoic rocks onshore norway and sweden. in the oslo graben, bitumen has been found in fractures in ordovician limestones and corals, and in southern norway, close to the oslo graben, oil and gas are trapped in permian volcanic intrusive rocks (pedersen et al. 2005, 2007). in southern sweden, at österplana, oil has been found in upper ordovician carbonate rocks (pedersen et al. 2007). pedersen et al. (2005, 2007) suggested that the petroleum was generated from lower palaeozoic source rocks. lower palaeozoic source rocks and oils are known from the baltic area, where lower silurian shales constitute the principal source (zdanaviciuté & bojesen-koefoed 1997). moreover, oil has been generated from palaeozoic units in sweden (sivhed et al. 2004 and references therein). the up to 90 m thick middle cambrian – lower ordovician alum shale, which is exposed in skåne, on bornholm and in southern norway, possesses source-rock properties. the black marine shales are highly organic-rich with an organic matter content locally up to 30% (thomsen et al. 1987; bharati et al. 1992). in central sweden, the alum shale is immature to marginally mature and has hi values up to above 600 mg hc/g toc, but in the fennoscandian border zone the shales are overmature (buchardt et al. 1986; buchardt & lewan 1990; bharati et al. 1992). the alum shale was encountered in the slagelse-1 and terne-1 wells. in the slagelse-1 well, toc contents are below the reliable detection limit of the carbon analyser, but in the terne-1 well, toc contents range from 5.39–10.08 wt%. the alum shale in the terne-1 and slagelse-1 wells is, however, overmature, with hi values <8 mg hc/g toc in the terne-1 well. high toc contents and hi values were recorded in the upper silurian nøvling formation in the rønde-1 well, where the formation consists of interbedded basalts, grey and red-brown claystones and sandstones with some carbonates (christensen 1971, 1973). the samples are, however, contaminated by various drilling mud additives, such as diesel, black magic®, and starch-based mud, which render the data unreliable. upper carboniferous strata were encountered by the hans-1 well, but no source-rock data are available. the drilled redbed section consists of interbedded sandstones, siltstones and claystones overlain by volcanic rocks and claystones, which are regarded to have no source-rock potential based on the lithology (michelsen & nielsen 1991, 1993). upper carboniferous coals are, however, significant gas source rocks regionally in northwest europe and the southern north sea (lokhorst 1998; gautier 2003). west of the study area, the carboniferous has been drilled by nine released wells in the danish central graben and the southern part of the norwegian and uk central graben; these strata are of early carboniferous age (bruce & stemmerik 2003). thin lower carboniferous coals were encountered in the gert-2 well, and these coals possess a gas/condensate generation potential inherited from the distribution and thermal maturity of potential source rocks 37 original vegetation, which is a general aspect of car boniferous coals (petersen 2006; petersen & nytoft 2006, 2007a, b). reworked carboniferous palynomorphs are common in the jurassic of the norwegian–danish basin suggesting that carboniferous strata, probably coal-bearing, were originally more widespread (nielsen & koppelhus 1990) and may be preserved below the regional permian unconformity in local deep grabens. these strata may thus potentially constitute a deep-seated source for gas/con densate. in summary, the well data described here do not demonstrate the occurrence of viable source rocks in the pre-per mian succession. it should be acknowledged, however, that the silurian is only known from the rønde-1, nøvling-1 and terne-1 wells where less than 450 m of the up to 2600 m thick succession have been investigated (christensen 1971, 1973; michelsen & nielsen 1991, 1993). similarly, the distribution and nature of carboniferous strata are very poorly known and it can be speculated that gas/condensateprone carboniferous strata may be preserved in local grabens and half-grabens, as recorded from the central graben (bruce & stemmerik 2003). permian units strata of zechstein and rotliegend age are present in nine wells, and results of source-rock analyses are available from five of them (c-1, nøvling-1, rønde-1, slagelse-1, sæby-1; fig. 1). no petroleum generation has been observed in the analysed well sections. high toc contents and hi values in the rotliegend group in the rønde-1 well, where the formation is mainly composed of reddish brown sandstones (jacobsen 1971; nielsen & japsen 1991), are artefacts from drilling mud additives. cuttings samples, showing a resemblance to black shale, from the rotliegend section (5092–5260 m) in the feli cia-1a well, yield extraordinarily high hi values (exceeding 1000 mg hc/g toc) and low tmax from 425–440ºc (petersen et al. 2003b). however, these data are flawed as the samples are contaminated by oil-based drilling mud (shellsol d-70), which was used below a depth of about 2000 m. core samples collected from thin black, shaly intervals yield hi values <32 mg hc/g toc and tmax values from 480–494ºc. this indicates overmaturity of the organic matter and no source potential. visual inspection of the dispersed organic matter (dom) by reflected light microscopy (white and fluorescing-inducing blue light) reveals lack of fluorescence of the dom that is classified as vitrinite (type iii kerogen) and inertinite (type iv kerogen). lack of fluorescence and visible liptinitic material is in agreement with the high organic maturity. tmax values of 480–494ºc correspond roughly to a vitrinite reflectance range of 1.6–1.9%ro, and at this maturity level the fluorescence behaviour of all types of organic matter has disappeared. on the basis of presently available data, therefore, the permian succession does not exhibit petroleum generation potential, although it is acknowledged that data are few and the permian potential cannot be categorically discounted. thin black, bituminous shales possibly equivalent to the kupferschiefer in the north german basin were encountered in the rønde-1 well (jacobsen 1971) and thin shales appear to be locally present in the zechstein succession. the distribution of these facies in the norwegian–danish basin is poorly known. triassic units triassic strata are dominated by continental to marginal marine sandstones, mudstones, marls, carbonates and evaporites, and good quality source rocks are not common in the triassic. petroleum generation potential has been detected in upper triassic strata in the hans-1, mejrup-1, rønde-1 and skagen-2 wells. a few hi values up to above 500 mg hc/g toc have been recorded from mudstones of the marine oddesund formation in the rønde-1 well, and in the mejrup-1 well a c. 8 m thick interval in the brackish vinding formation has hi values up to nearly 700 mg hc/g toc. in the hans-1 well, where parts of the gassum formation consist of aggrading parasequences of coastal plain deposits with coal beds, a c. 10 m thick interval possesses gas generation potential. in the skagen-2 well, the uppermost part of the gassum formation and the lowermost part of the f-i member of the fjerritslev formation comprise lagoonal sediments with a restricted capacity to generate liquid petroleum (see below). in the mejrup-1 well, the gassum formation is more fine-grained than normal, and certain mudstone intervals have hi values up to 534 mg hc/g toc; the average hi value for a c. 73 m thick mudstonedominated section is 216 mg hc/g toc. this unusual development of the formation in mejrup-1 is probably related to the position of the well in the secondary rim syncline of the vejrum salt dome, in the centre of the basin. movement of salt influenced depositional patterns at many locations in the basin, as shown by seismic data, but only few wells have drilled rim synclines and none show a development similar to that of the gassum formation in the mejrup-1 well. however, detailed seismic mapping of the upper triassic may determine if similar depositional situations and sufficient burial of upper triassic potential 38 source rocks occur adjacent to other salt structures. in summary, apart from a few local occurrences of upper triassic units with a limited potential, the triassic does not possess a petroleum generation potential. lower jurassic units with the sole exception of the c-1 well, the lower jurassic offshore marine mudstones of the fjerritslev formation have been encountered in all the investigated wells. all four members of the fjerritslev formation (f-i – f-iv) are, however, not present in all wells due to middle jurassic uplift and erosion over much of the basin (andsbjerg et al. 2001; nielsen 2003). as a result of the uplift and associated erosion, the f-i member at the base of the formation is regionally the most widespread, being present in all but one well, whereas the f-iv member in the uppermost part of the formation is only present in 20 of the 33 wells. the f-ii and f-iii members are present in 25 of the wells. 400 450 500 0 100 200 300 400 500 600 700 f-1 farsø-1 felicia-1 fjerritslev-2 hans-1 hobro-1 hyllebjerg-1 inez-1 j-1 k-1 kvols-1 lavø-1 mejrup-1 mors-1 nøvling-1 rødding-1 rønde-1 skagen-2 skive-1 slagelse-1 sæby-1 terne-1 vedsted-1 vinding-1 voldum-1 års-1 h yd ro ge n in de x (m g h c /g t o c ) tmax (°c) type ii type iii fig. 23. hydrogen index versus tmax plot of 497 samples from the f-i member of the lower jurassic fjerritslev formation. the f-i member is a poor source rock, generally with hi values <100 mg hc/g toc. 39 f-i member only in the skagen-2 well does the f-i member possess a potential for petroleum generation (fig. 23). as noted above (triassic units), a 10 m thick interval spanning the uppermost part of the gassum formation and the lowermost part of the f-i member shows hi values ranging from 255–273 mg hc/g toc. this interval is interpreted to consist of stacked lagoonal deposits that are only locally developed (sequence fj 9 of nielsen 2003). f-ii member a marginal petroleum generation potential has been recorded in the f-ii member in the farsø-1 and mors-1 wells, whereas the member in the hobro-1 well shows a somewhat better potential (fig. 24), albeit over a very narrow interval. within a 6 m thick interval in hobro-1, the hi values range from 203–340 mg hc/g toc. the topmost 10 m of the f-ii member in mors-1 has an average hi value of 183 mg hc/g toc, whereas the average hi value of the member in farsø-1 is 156 mg hc/g toc. the f-ii mem400 450 500 0 100 200 300 400 500 600 700 f-1 farsø-1 felicia-1 fjerritslev-2 flyvbjerg-1 frederikshavn-1 frederikshavn-2 hans-1 hobro-1 hyllebjerg-1 j-1 k-1 mejrup-1 mors-1 nøvling-1 rødding-1 rønde-1 skive-1 sæby-1 terne-1 vedsted-1 voldum-1 års-1 h yd ro ge n in de x (m g h c /g t o c ) tmax (°c) type ii type iii fig. 24. hydrogen index versus tmax plot of 176 samples from the f-ii member of the lower jurassic fjerritslev formation. the f-ii member is a poor source rock, generally with hi values <150 mg hc/g toc. 40 ber in the latter two wells can thus principally be regarded as gas-prone. in the other wells, the mudstones of the f-ii member possess no source-rock potential. generally, the source-rock potential of the f-ii member can thus be regarded as limited and primarily gas-prone. f-iii and f-iv members the distribution of the f-iii and f-iv members is controlled by post-depositional erosion related to the regional early middle jurassic uplift that influenced most of the norwegian–danish basin and the fennoscandian border zone. the map in fig. 15 displays the area within which sediments of f-iii or f-iv are preserved; in some wells the combined thickness of the f-iii and f-iv members amounts to nearly 400 m (fig. 25a; see also table 2). the isopach map suggests that the largest combined thicknesses occur in the himmerland graben (fig. 25b). the entire stratigraphic interval is only preserved in the sorgenfrei–tornquist zone. the analyses of the 33 well sections show that the average quality of the f-iii and f-iv members, in terms of potential source rocks for oil generation, is highly variable both stratigraphically and geographically (figs 26, 27). in the kvols-1 and rønde-1 wells, part of the f-iii member constitutes an excellent potential oil source rock (fig. 26). the uppermost alginite-bearing 40 m of the member in kvols-1 has an average toc content of 2.95 wt%, and the interval displays hi values consistently >300 mg hc/g toc, with a maximum value of 529 mg hc/g toc and an average value of 429 mg hc/g toc (fig. 28, 29a). the marine mudstones in this interval contain abundant algal-derived organic material composed of fluorescing, amorphous organic matter (aom) and alginite of the tasmanites and leiosphaeridia types (fig. 28). associated framboidal pyrite testifies to the oxygen-deficient, organic-rich conditions in the sediment during deposition. this highly oil-prone section overlies a c. 25 m thick interval with an average hi value of 243 mg hc/g toc. in rønde-1, the topmost c. 18 m of the f-iii member are highly oil-prone, with hi values ranging from 263–428 mg hc/g toc, averaging 355 mg hc/g toc (fig. 29b). similarly, the f-iii member in the haldager-1 well contains a c. 25 m thick oil-prone interval, with hi values reaching 425 mg hc/g toc and averaging 323 mg hc/g toc (fig. 30a). in a number of other wells, the f-iii member also posses ses a variable petroleum generation potential. the farsø-1 and j-1 wells show maximum hi values around 300 mg hc/g toc and average hi values of 214 mg hc/g toc and 202 mg hc/g toc, respectively, indicating a limited liquid petroleum generation potential (fig. 30b). the mors-1 well has an average hi value of the same order, but the maximum hi value only reaches 266 mg hc/g toc (fig. 31a). in the hobro-1 well, a c. 12 m thick interval of the f-iii member has an average hi of 233 mg hc/g toc (fig. 31b). the f-iii member in fjerritslev-2, hyllebjerg-1 (fig. 32a) and voldum-1 (fig. 32b) contains intervals with average hi values around 190 mg hc/g toc, whereas in the års-1 well, the average value is only 156 mg hc/g toc. the f-iii member of these latter four wells is considered to be principally gas-prone. in the rønde-1 well, the f-iv member (55 m thick) is an excellent, organic-rich (average toc of 3.58 wt%), oilprone source rock with an average hi value of 435 mg hc/g toc and a maximum hi value of 543 mg hc/g toc (figs 27, 29b). the organic material in this highly oil-prone interval consists of abundant fluorescing aom, detrital liptinite and leiosphaeridia type alginite (fig. 33). both the aom and detrital liptinite are probably composed of degraded algal material. the occurrence of framboidal pyrite is indicative of oxygen-deficient conditions during deposition of the organic-rich shales. a nearly 50 m thick section of the f-iv member in the hyllebjerg-1 well has an average hi value of 210 mg hc/g toc, with a maximum hi value of 371 mg hc/g toc (fig. 32a). table 2. net source-rock (sr) thicknesses, f-iii and f-iv members of the fjerritslev formation well f-iii net sr f-iii gross f-iii f-iv net sr f-iv gross f-iv sum net sr sum gross sum average hi*, net sr interval (m) (m) net/gross (m) (m) net/gross (m) (m) net/gross (number of samples) haldager-1 31 117 0.26 9 127 0.07 40 244 0.16 294 (7) hobro-1 19 154 0.12 7 32 0.22 26 186 0.14 229 (4) hyllebjerg-1 35 228 0.15 21 55 0.38 56 283 0.20 246 (14) j-1 19 98 0.19 0 43 0 19 141 0.13 267 (2) kvols-1 79 201 0.39 4 18 0.22 83 219 0.38 369 (28) mors-1 118 147 0.80 5 25 0.20 123 172 0.72 221 (17) rønde-1 48 145 0.33 44 55 0.80 92 200 0.46 404 (12) skagen-2 7 59 0.12 0 29 0 7 88 0.08 203 (1) sæby-1 0 53 0 4 38 0.11 4 91 0.04 261 (1) voldum-1 14 111 0.13 0 0 0 14 111 0.13 215 (2) * average hi in net sr intervals based on hi values ≥200 mg hc/g toc 41 f-iv f-iii 50 km 400 m 200 m 0 m high fault national border a jylland 50 km limit of f-iii and f-iv members >400 m 300–400 m 250–300 m 200–250 m 150–200 m <150 m high fault national border jylland b fig. 25. a: thickness of the lower jurassic lithostratigraphic f-iii and f-iv members in the fjerritslev formation in the studied wells; the base of each schematic stratigraphic column is located at the well site. b: isopach map of the combined thickness of the f-iii and f-iv members. the largest thicknesses occur in the himmerland graben. 42 a marginal gas generation potential is shown by the member in the j-1 well (fig. 30b), where the average hi value is only 130 mg hc/g toc and the maximum hi value is 178 mg hc/g toc. the average hi values in the f-iii and f-iv members, calculated from the entire thickness of the members, classify the majority of well-sections as gas-prone and a few as gas-/oil-prone (fig. 34). the source-rock quality of the f-iii and f-iv members, as shown above, varies considerably between wells and within the members in individual wells; average values thus mask specific oil-prone intervals. to illustrate further the evaluation of the source rocks in the f-iii and f-iv members, a ‘net quality map’ was constructed showing the cumulative thickness of intervals with hi values >200 mg hc/g toc (fig. 35a; table 2) and the net/gross ratio (cumulative thickness/total thickness of f-iii + f-iv members). the map thus displays the occurrence and thickness of source rocks with a mixed gas/oil or oil generation potential and suggests that the thickest cumulative source-rock section with hi >200 mg hc/g toc occurs in the basin centre (kvols-1, mors-1 and rønde-1 wells), where the best source-rock quality also occurs. the kvols-1 and rønde-1 wells contain about 80–90 m net oil-prone source rock with average hi values of 369 and 404 mg hc/g toc, respectively. it is also notable that the haldager-1 well in the sorgenfrei–tornquist zone contains c. 40 m net source rock with an average hi of 294 mg hc/g toc. the thinner net source-rock thicknesses 400 450 500 0 100 200 300 400 500 600 700 børglum-1 f-1 farsø-1 felicia-1 fjerritslev-2 flyvbjerg-1 frederikshavn-1 frederikshavn-2 haldager-1 hobro-1 hyllebjerg-1 j-1 k-1 kvols-1 mejrup-1 mors-1 rødding-1 rønde-1 skive-1 sæby-1 terne-1 vedsted-1 voldum-1 års-1 h yd ro ge n in de x (m g h c /g t o c ) tmax (°c) type ii type iii fig. 26. hydrogen index versus tmax plot of 364 samples from the toarcian f-iii member of the fjerritslev formation. the f-iii member is principally a good, mixed gas/oil-prone source rock, but with some wells containing highly oil-prone intervals with hi values >350 mg hc/g toc. 43 in the hobro-1, hyllebjerg-1 and voldum-1 wells, which are also situated centrally in the basin, may partly be explained by erosion of the upper part of the f-iv member during the middle jurassic uplift event. in other centrally placed wells, such as mejrup-1, rødding-1 and skive-1, the entire f-iv member and most of the f-iii member have been eroded, resulting in potential source rocks being thin or absent in these areas. the limiting factor for the petroleum potential of the study area is the thermal maturity of the richest source-rock units, the f-iii and f-iv members. it has been shown that the depth to the top of the oil window, corresponding to a vr of 0.6%ro, is about 3050–3100 m (figs 20a, 22a). using this threshold, the f-iii and f-iv members are thermally immature in the investigated wells, also demonstrated by the tmax values (fig. 35b). the largest discrepancy between tmax and vr values is observed in the terne-1 well, which compared to the other wells also yields unusually low vr values (figs 18, 19, 21). the notable lack of hydrocarbon shows in the wells drilled in the study area, including the central deep part, supports the suggestion that the f-iii and f-iv source rocks had not, prior to postearly cretaceous uplift, been buried to the necessary depth for petroleum generation to occur. the mors-1 and års-1 wells represent locations where the fjerritslev formation has been buried below the depth of the top of the oil window, but in both wells the petroleum generation potential of the sediments is relatively poor to non-existent. 400 450 500 0 100 200 300 400 500 600 700 børglum-1 farsø-1 fjerritslev-2 flyvbjerg-1 frederikshavn-2 gassum-1 haldager-1 hobro-1 hyllebjerg-1 j-1 k-1 kvols-1 mors-1 rønde-1 skagen-2 sæby-1 terne-1 vedsted-1 års-1 h yd ro ge n in de x (m g h c /g t o c ) tmax (°c) type ii type iii fig. 27. hydrogen index versus tmax plot of 142 samples from the toarcian f-iv member of the fjerritslev formation. the f-iv member is principally a good, mixed gas/oil-prone source rock, but with the hyllebjerg-1 and rønde-1 wells containing considerably richer intervals with hi values reaching c. 450–550 mg hc/g toc in rønde-1. 44 c d p p i i aom t e f p p aom l l a b p p tl l aom fig. 28. paired photomicrographs (reflected light, oil immersion; scale bar is c. 30 µm) of the organic material in the oil-prone f-iii member in the kvols-1 well; left (a, c, e): white light; right (b, d, f ): fluorescence-inducing blue light. a-d: fluorescing, amorphous organic matter (aom; probably algal-derived) and alginite with tasmanites (t) and leiosphaeridia (l) morphology in cuttings from c. 1996 m. inertinite (i) and framboidal pyrite (p) are also present. toc = 2.56 wt%, hi = 416 mg hc/g toc. e, f: fluorescing aom (probably algal-derived) and several alginites with leiosphaeridia (l) morphology in cuttings from c. 2012 m. abundant framboidal pyrite (p) present. toc = 3.57 wt%, hi = 487 mg hc/g toc. 45 d ep th ( m ) bø. fm h. s. fm fj. fm, f-iv fj. fm, f-iii fj. fm, f-ii mfs-15 ts-15 sb-15 mfs-14 ts-14 sb-14 0 40 80 120 160 120 80 40 0 200 400300 600 0.1 1 10 100 gamma ray (api) kvols-1 dt (µsec./ft) hi (mg hc/g toc) s2 (mg hc/g rock) 2200 2000 2150 2100 2050 1950 1959 1974 1992 2006 2062 2068 2094 2183 2192 1931 po or fa ir g oo d v er y go od ex ce lle nt g as g as / oi l o il a bø. fm fj. fm, f-iv fj. fm, f-iii mfs-15 ts-15 sb-15 mfs-14 ts-14 2100 2138 2193 2205 2240 2257 2285 2338 2342 sb-14 0 40 80 120 0 200 400300 600 0.1 1 10 100 gamma ray (api) hi (mg hc/g toc) s2 (mg hc/g rock) fa ir ex ce lle nt g as o il 400 300 200 40100 dt (µsec./ft) 2100 2150 2200 2250 2300 2350 d ep th ( m ) po or g oo d v er y go od g as / oi l rønde-1 b fig. 29. plots showing well logs, hydrogen index (hi) values, s2 yields, sequence stratigraphic key surfaces and formations in the kvols-1 (a) and rønde-1 (b) wells. note the oil-prone upper part of the f-iii member in kvols-1 and the oil-prone f-iv member in rønde-1. dt, sonic velocity; mfs, maximum marine flooding surface; sb, sequence boundary; ts, transgressive surface. bø., børglum; fj., fjerritslev; h.s., haldager sand. 46 0 40 80 120 d ep th ( m ) -100 -50 0 50 100 150 0 200 400300 600 0.1 1 10 100 1284 1280 1295 1324 1350 1372 1403 1407 1471 1485 1495 1524 1400 1500 1450 1350 1300 1250 mfs-18 sb-18 mfs-17 sb-17 mfs-16 sb-16 mfs-15 ts-15 sb-15 td fj. fm, f-iv fj. fm, f-iii gamma ray (api) sp (mv) hi (mg hc/g toc) s2 (mg hc/g rock) g as g as / oi l o il po or fa ir g oo d v er y go od ex ce lle nt haldager-1 a h. s. fm fj. fm, f-ivsb-16 fj. fm, f-iii mfs-15 ts-15 sb-15 mfs-14 fj. fm, f-ii 1092 1111 1132 1154 1190 1208 1222 1237 1252 1264 sb-14 0 40 80 120 0 200 400300 600 0.1 1 10 100 gamma ray (api) hi (mg hc/g toc) s2 (mg hc/g rock) po or fa ir g oo d v er y go od ex ce lle nt g as g as / oi l o il 160 120 80 40 dt (µsec./ft) 1100 1150 1200 1250 1300 d ep th ( m ) j-1 b fig. 30. plots showing well logs, hydrogen index (hi) values, s2 yields, sequence stratigraphic key surfaces and formations in the haldager-1 (a) and j-1 (b) wells. note the oil-prone interval in the f-iii member in haldager-1. dt, sonic velocity; sp, self-potential; td, total depth. for further abbreviations, see fig. 29. 47 fj. fm, f-iv fj. fm, f-iii sb-16 mfs-15 ts-15 sb-15 mfs-14 ts-14 2155 2179 2213 2240 2252 2275 2327 2335 sb-14 0 40 80 120 0 200 400300 600 0.1 1 10 100 gamma ray (api) hi (mg hc/g toc) s2 (mg hc/g rock) po or fa ir g oo d v er y go od ex ce lle nt g as g as / oi l o il 160 120 80 40 dt (µsec./ft) 2150 2200 2250 2300 2350 d ep th ( m ) mors-1 a h. s. fm fj. fm, f-iv mfs-16 fj. fm, f-iii mfs-15 ts-15 sb-15 mfs-14 fj. fm, f-iisb-14 1884 1923 1932 1955 1972 2008 2013 2039 2109 2122 2140 fj. fm, f-i 0 40 80 120 0 200 400300 600 0.1 1 10 100 gamma ray (api) hi (mg hc/g toc) s2 (mg hc/g rock) o il 160 120 80 40 dt (µsec./ft) 1900 1950 2000 2050 2100 2150 d ep th ( m ) fig34 po or fa ir g oo d v er y go od ex ce lle nt g as g as / oi l hobro-1 b fig. 31. plots showing well logs, hydrogen index (hi) values, s2 yields, sequence stratigraphic key surfaces and formations in the mors-1 (a) and hobro-1 (b) wells. note the c. 125 m thick interval in the f-iii member in mors-1 in which hi values exceed 200 mg hc/g toc. dt, sonic velocity. for further abbreviations, see fig. 29. 48 1894 1913 1922 1927 mfs-16 1947 sb-16 1977 2060 mfs-15 2093 ts-15 2114 sb-15 2145 mfs-14 2205 2213 sb-14 0 40 80 120 0 200 400300 600 0.1 1 10 100 gamma ray (api) hi (mg hc/g toc) s2 (mg hc/g rock) po or fa ir g oo d g as g as / oi l o il 160 120 80 40 dt (µsec./ft) 1900 1950 2000 2050 2100 2150 2200 2250 d ep th ( m ) fl. fm h. s. fm fj. fm, f-ii fj. fm, f-iii fj. fm, f-iv fig35 v er y go od ex ce lle nt hyllebjerg-1 a 1500 1400 1550 1450 1350 bø. fm 1393 h. s. fm 1423 fj. fm, f-iii 1533 fj. fm, f-ii 1562 fj. fm, f-i 0 40 80 120 0 200 400300 600 0.1 1 10 100 gamma ray (api) hi (mg hc/g toc) s2 (mg hc/g rock) po or fa ir g oo d v er y go od ex ce lle nt g as g as / oi l o il 160 120 80 40 dt (µsec./ft) d ep th ( m ) 1379 fig_36 voldum-1 b fig. 32. plots showing well logs, hydrogen index (hi) values, s2 yields, sequence stratigraphic key surfaces and formations in the hyllebjerg-1 (a) and voldum-1 (b) wells. in hyllebjerg-1, note the increased hi values in the upper part of the f-iii member and in the middle part of the f-iv member. dt, sonic velocity. for further abbreviations, see fig. 29. 49 middle jurassic units in the yme field in the egersund basin, where the bryne formation (equivalent to the haldager sand formation) is thickly developed, the formation contains shales that show a good to excellent oil generation potential with toc values of 2–13 wt% and hi values of 100–480 mg hc/g toc. coal seams within the middle jurassic – lower upper jurassic of this field are also oil-prone in places, with hi values exceeding 400 mg hc/g toc. similarly, middle jurassic coals and carbonaceous lacustrine–brackish shales of the lulu formation are considered to have sourced the oil and gas/condensate accumulations in the lulita and harald fields in the søgne basin of the north sea. in this area, the type of generated petroleum is considered to have been controlled by lateral coal facies variations related to the proximity of peat formation to the coeval coastline (petersen et al. 1998, 2000; petersen & brekke 2001). coals formed in the coastal reaches of the mires are more oil-prone than their more landward equivalents. in the study area, the middle jurassic haldager sand formation is dominated by fluvial, estuarine and shallow marine sandstones interbedded with marine and lacustrine mudstones and thin coal seams (nielsen 2003). the generation potential of the formation is low, with only a few exceptions. in the haldager-1 well, two samples from a 4–5 m thick marine mudstone have high hi values (342 and 637 mg hc/g toc), and the terne-1 well shows hi values of 260 mg hc/g toc. both wells are situated in the sorgenfrei–tornquist zone where the middle jurassic is thickest. there may thus be a relationship between the gross thickness of the middle jurassic succession, the palaeoc d p p aom l l t a b pp pp p p l aom fig. 33. paired photomicrographs (reflected light, oil immersion; scale bar is c. 30 µm) of the organic material in cuttings sample (c. 2152 m) from the oil-prone f-iv member in the rønde-1 well; left (a, c): white light; right (b, d): fluorescence-inducing blue light. the cuttings contain an abundance of fluorescing, amorphous organic matter (aom) and detrital liptinite (probably algal-derived) together with leiosphaeridia (l) and tasmanites (t) alginite. p, framboidal pyrite. toc = 5.81 wt%, hi = 543 mg hc/g toc. 50 3.38 1.66 0.97 0.75 0.07 0.25 1.32 1.10 0.15 0.59 0.92 1.02 1.68 1.43 1.69 0.77 1.01 1.95 0.15 0.75 2.48 1.7 0.18 0.690.31 0.57 0.44 0.65 2.68 1.38 0.90 1.05 0.92 1.18 1.27 2.54 2.63 1.29 1.28 1.14 2.74 1.30 4.42 1.92 5.2 1.68 0.72 0.89 0.61 1.17 0.62 1.04 fault national border high toc (wt%) 50 km 0.75 0.75 oil: hi = 300–600 mg hc/g toc gas/oil: hi = 200–300 mg hc/g toc gas: hi <200 mg hc/g toc s2 (mg hc/g rock)a jylland 1.49 3.1 0.35 3.53 0.170.161.1 2.03 1.19 1.04 1.22 1.02 1.38 1.28 16.09 0.5 2.47 2.16 0.93 3.12 0.93 0.290.481.42 1.4 1.38 1.37 1.04 1.68 1.03 1.47 3.58 0.6 1.2 4.07 fault national border high toc (wt%) 50 km 0.34 0.34 oil: hi = 300–600 mg hc/g toc gas/oil: hi = 200–300 mg hc/g toc gas: hi <200 mg hc/g toc s2 (mg hc/g rock) jylland b fig. 34. the average source-rock quality of the lower jurassic f-iii member (a) and f-iv member (b) of the fjerritslev formation. 51 0.16 0.20 123 0.13 40 56 0.14 19 7 4 50 km 0.08 net/gross ratio net thickness in metres average hi of samples with hi >200 (no. of samples) 7 fig. 40 skagen-2 203 (1) sæby-1 261 (1) haldager-1 294 (7) j-1 267 (2) mors-1 221 (17) hobro-1 229 (4)kvols-1 369 (28) rønde-1 404 (12) 0.13 0.46 14 92 hyldebjerg-1 246 (14) 203 (1) 0.08 0.04 0.72 26 0.38 83 voldum-1 215 (2) fault national border high oil: hi = 300–600 mg hc/g toc gas/oil: hi = 200–300 mg hc/g toc gas: hi <200 mg hc/g toc a jylland 0.40 0.46 0.49 0.45 0.45 0.52 0.51 0.56 0.58 0.51 0.53 0.50 0.54 0.34 fault national border high 50 km immature: <435°c early oil window: 435–445°c tmax vitrinite reflectance (%ro) fig. 41 0.34 b jylland fig. 35. a: cumulative net source-rock thickness of the intervals in the f-iii and f-iv members with hi values exceeding 200 mg hc/g toc, i.e. a source rock with a mixed gas/oil or oil generation potential. b: thermal maturity of the f-iii and f-iv members of the fjerritslev formation. vitrinite reflectance and tmax values indicate immaturity with regard to petroleum generation. 52 geographic position and the occurrence of shales and coals with an oil generation potential. oil-prone coals in the middle jurassic may be best developed in areas with most pronounced subsidence and hence relatively large rates of formation of accommodation space during deposition. upper jurassic – lower cretaceous units in most wells, the uppermost jurassic – lowermost creta ceous frederikshavn formation is dominated by shallow marine and paralic siltstones and sandstones (michelsen et al. 2003) with no petroleum generation potential. the average hi of the entire frederikshavn formation indicates a gas-prone source potential (fig. 36), but the formation contains intervals with good to excellent oil generation po tential, as demonstrated by the gassum-1, hyllebjerg-1, skagen-2, sæby-1, terne-1 and voldum-1 wells (figs 37, 38a; table 3). the section in terne-1 is particularly noteworthy, with hi values >1100 mg hc/g toc, although these values in part reflect contamination by gel mud and cement applied during drilling. nevertheless, solventextracted samples are still encouraging, with a cumulative net source-rock unit of c. 150 m containing on average 5.7 wt% toc and hi values of the extracted samples reaching 580 mg hc/g toc, averaging 478 mg hc/g toc (fig. 38a; table 3). microscopical kerogen analyses show an abundance of amorphous algal organic matter in the form of filamentous lamalginite and alginite with morphology similar to the extant fresh to brackish water botryococcus algae (type i kerogen; fig. 39). brackish, oxygen-deficient conditions during deposition are suggested by the abundance of framboidal pyrite associated with the organic matter (fig. 39). the probable lacustrine origin of these deposits in terne-1 suggests a local development, as the formation is typically of shallow marine to offshore origin (michelsen et al. 2003). the skagen-2 well contains a c. 78 m thick net sourcerock interval with an average hi value of 241 mg hc/g 50 km fault national border high toc (wt%)0.34 0.34 0.52 0.71 0.72 0.86 0.98 0.910.56 0.90 0.26 1.18 0.88 0.95 0.45 0.61 0.53 0.26 1.09 1.74 2.08 1.07 0.55 0.23 0.37 0.07 0.50 0.27 0.53 0.26 0.20 0.47 0.76 0.54 4.39 23.78 0.21 0.45 0.11 0.52 0.24 0.94 0.49 0.73 0.32 0.20 0.17 oil: hi = 300–600 mg hc/g toc gas/oil: hi = 200–300 mg hc/g toc gas: hi <200 mg hc/g toc s2 (mg hc/g rock) 2.2 1.01 jylland fig. 36. average source-rock quality of the entire uppermost jurassic – lowermost cretaceous frederikshavn formation. well net sr gross net/ average hi*, net sr interval (m) (m) gross (number of samples) gassum-1 17 101 0.17 320 (3) hyllebjerg-1 29 146 0.20 243 (4) skagen-2 78 176 0.44 241 (7) sæby-1 10 105 0.10 329 (1) terne-1 150 258 0.58 478 (10) voldum-1 26 66 0.39 257 (3) * average hi in net sr intervals based on hi values ≥200 mg hc/g toc table 3. net source-rock (sr) thicknesses, frederikshavn fm 53 toc and a maximum value close to 350 (fig. 38a). in the gassum-1 well, most hi values are low but a few samples yield hi values >300 mg hc/g toc (fig. 38a); these high values are abnormal compared to the general trend and may reflect thin layers with higher quality kerogen in an otherwise sand-dominated succession. the hyllebjerg-1 well has a c. 29 m thick net source-rock section with an average hi value of 243 mg hc/g toc, whereas the average hi value of c. 26 m net source rock in the voldum-1 well is 257 mg hc/g toc (fig. 38a; table 3). the frederikshavn formation thus locally contains good to excellent oil source rocks in the study area. towards the west in the egersund basin, the broadly time-equivalent tau formation is known as the principal source for oil, and in the north sea the uppermost jurassic – lowermost creta ceous marine shales of the farsund formation and equivalents (kimmeridge clay, mandal and draupne formations) are well known as the primary oil source rocks (e.g. ineson et al. 2003). as for the lower jurassic fjerritslev formation, the major problem in the danish area is to find areas where the frederikshavn succession is sufficiently buried to be thermally mature with regard to petroleum generation. vr values from 0.36–0.53% ro and the majority of tmax values <430–435°c indicate that the potential source rocks are thermally immature (fig. 38b). 400 450 500 550 h yd ro ge n in de x (m g h c /g t o c ) gassum-1 (n = 15) hyllebjerg-1 (n = 17) skagen-2 (n = 15) terne-1 (n = 14, extracted) voldum-1 (n = 7) 0 100 200 300 400 500 600 700 800 900 1000 tmax (°c) type i type ii type iii fig. 37. hydrogen index versus tmax plot of 68 samples from the uppermost jurassic – lowermost cretaceous frederikshavn formation. in the terne-1 well, in particular, the frederikshavn formation constitutes an excellent oilprone source rock with hi values exceeding 500 mg hc/g toc. 54 fig . 44 skagen-2 241 (7) sæby-1 329 (1) gassum-1 320 (3) voldum-1 257 (3) terne-1 478 (10) hyldebjerg-1 243 (3) gas/oil: hi = 200–300 mg hc/g toc fault national border high oil: hi = 300–600 mg hc/g toc 50 km 0.4 net/gross ratio net thickness in metres average hi of samples with hi >200 (no. of samples) 78 78 0.4 10 29 17 26 0.39 0.58 150 0.17 0.2 0.1 241 (7)a jylland gas: hi <200 mg hc/g toc fault national border high immature: <435°c : early oil window: 435–445°c 0.34 0.34 0.41 0.44 0.44 0.45 0.48 0.51 0.46 0.46 0.44 0.51 0.53 50 km tmax vitrinite reflectance (%ro) b jylland fig. 38. a: cumulative net source-rock thickness of the intervals in the frederikshavn formation with hi values exceeding 200 mg hc/g toc, i.e. a source rock with a mixed gas/oil or oil generation potential. b: thermal maturity of the frederikshavn formation. vitrinite reflectance values indicate immaturity, whereas tmax values may suggest early oil window maturity in four wells located in the central part of the basin or in the sorgenfrei–tornquist zone. 55 c d b p p b b fl e f p b bl b a b p fl fl b fig. 39. paired photomicrographs (reflected light, oil immersion; scale bar is c. 30 µm) of the organic material in the oil-prone lacustrine frederikshavn formation in the terne-1 well; left (a, c, e) white light; right (b, d, f ): fluorescence-inducing blue light. a-f: abundance of fluorescing, filamentous lamalginite (fl) and botryococcus-type alginites (b) of varying size in cuttings from c. 200–210 m. abundant framboidal pyrite (p) has been formed within the large botryococcus-type alginite bodies. toc = 7.30 wt%, hiextracted = 498 mg hc/g toc. 56 lower cretaceous units in general, the lower cretaceous succession of the study area contains few and relatively thin potential oil source rocks. three wells, the lavø-1, sæby-1 and års-1, contain intervals in the lower cretaceous (vedsted formation or undifferentiated lower cretaceous) with a petroleum generation potential. in the sæby-1 well, a c. 20 m thick organic-rich interval has toc contents up to 5.28 wt% and hi values from 320–472 mg hc/g toc, averaging 388 mg hc/g toc; the interval is thus an excellent potential oil and gas source rock. similarly, a c. 20 m thick interval in the års-1 well possesses a mixed oil/gas generation potential, although the potential is poorer than in sæby-1 as the hi values range from 210–411 mg hc/g toc, averaging 294 mg hc/g toc. in the lavø-1 well, a c.18 m thick potential source-rock interval with hi values from 175–242 mg hc/g toc is present. the lower cretaceous is, however, thermally immature in all well sections. potential reservoirs the principal sedimentary units of interest with respect to potential reservoirs in the norwegian–danish basin are the sandstones of the upper triassic – lowermost jurassic gassum formation and the middle jurassic haldager sand formation. the growth of salt pillows caused local topographic relief that influenced the deposition of these two reservoir units, as well as the intervening fjerritslev for mation. the units thicken into rim synclines as indicated, for example, by the felicia-1/1a section, and may thin considerably over salt pillows. these principal potential reservoirs are reviewed below with respect to their gross distribution, thickness development and properties. a number of secondary potential reservoir units are also known from the norwegian–danish basin and the fennoscandian border zone, including the lower triassic bunter sandstone, the lower–upper triassic skagerrak formation, the lower jurassic f-ii member of the fjerritslev formation, the upper jurassic flyvbjerg formation and the uppermost jurassic – lowermost cretaceous frederikshavn formation. these secondary reservoir units are briefly described after the principal reservoirs. gassum reservoir shoreface and fluvial–estuarine sandstones interbedded with marine mudstones, lagoonal heteroliths and mudstones, lacustrine mudstones and thin coal seams occur in the gassum formation (nielsen 2003). in the himmerland graben, the sorgenfrei–tornquist zone and the skagerrak –kattegat platform, sandstones are commonly the dominant lithology, and petrophysical log evaluations typically show net-to-gross ratios of 0.3–0.7 and porosities of 15–25%. the formation is more sand-poor in the central part of the basin with net-to-gross ratios of 0.1–0.2 (e.g. mejrup-1, nøvling-1, vemb-1, vinding-1). the sandstones are predominantly well to moderately sorted, fineto medium-grained, locally coarse-grained and slightly pebbly. the shoreface sandstones occur as widespread sheets, 4–30 m thick, separated by marine transgressive mudstones and lagoonal heteroliths. thick fluvial–estuarine sandstones mostly overlie the major sb 5 sequence boundary (figs 10, 12; nielsen 2003). the formation is 50–150 m thick in central parts of the norwegian–danish basin, its thickness being influenced by proximity to salt structures and faults. the formation thickens to 170–200 m in the fault-bounded himmerland graben and the northern part of the sorgenfrei–tornquist zone. it thickens further to more than 300 m in the southern part of the fault zone where deposition of sand continued from the triassic until the early sinemurian (nielsen 2003). the thickness ranges from 69–205 m in the f-1, k-1, felicia-1/1a and j-1 wells (fig. 1). the large thickness (205 m) of the formation in felicia-1/1a, with thick mudstones in the middle part of the formation, probably reflects an excess of accommodation space in the rim syncline associated with the nearby large salt pillow. the thickness decreases to 10–80 m on the skagerrak–kattegat platform. it is generally assumed that the siliciclastic material was mainly supplied from the baltic shield to the north and east. however, the dominance of mineralogically mature and better sorted sandstones in the stenlille area and in the ullerslev-1 well suggests that sand may have been supplied from the erosion of older sediments, such as the triassic bunter sandstone on the ringkøbing–fyn high (larsen 1966; nielsen 2003). the gassum formation is utilised in geothermal energy installations onshore denmark at a depth of c. 1200 m (thisted, northern jylland) and is used for storage of natural gas in a structure at c. 1550 m depth in the eastern part of the basin (stenlille area; fig. 1). 57 haldager sand reservoir the haldager sand formation consists primarily of sandstones interbedded with thin mudstones. the sandstones are mediumto coarse-grained, slightly pebbly, commonly well to moderately sorted but locally poorly sorted. sandstones are the dominant lithology, and petrophysical log evaluations typically show net-to-gross ratios of 0.4–0.8 and porosities of 15–30%. in the sorgenfrei–tornquist zone, where subsidence continued despite regional uplift, the formation consists of four thick fluvial–estuarine to shallow marine sandstone units separated by marine and lagoonal–lacustrine mudstones (nielsen 2003). beyond the fault-bounded graben, in areas that experienced uplift in the early part of the middle jurassic, sandstones were mainly deposited by braided rivers, and the sand bodies are expected to be laterally coherent without significant primary hydraulic barriers. anomalies with respect to facies and thickness occur locally in rim synclines associated with salt structures. the distribution and thickness of the haldager sand reservoir are strongly influenced by regional syndepositional tectonism, local faulting and salt structures. sediments were supplied from the north and east, but deep erosion of triassic and older strata on the ringkøbing–fyn high and lower jurassic mudstones along the northern flank of the high added a substantial amount of material. as a result of the uplift of the ringkøbing–fyn high, high-energy braided rivers shed erosion products into the sorgen frei–tornquist zone, which experienced slow fault-controlled subsidence. between the fjerritslev and børglum faults, the formation attains a thickness of 30–175 m. out side the sorgenfrei–tornquist zone, the thickness and number of sandstone units decreases. on the skagerrak –kattegat platform, the formation is 15–50 m thick and in the central part of the basin it is 25–50 m thick. in the southern and south-western part of the study area, the formation is thin and has a patchy distribution with thicknesses below 10 m. additional reservoirs the lower triassic bunter sandstone and the lower–upper triassic skagerrak formations constitute additional potential reservoir units in the study area. the bunter sandstone formation consists of orange, red-brown and yellow-brown, mediumto fine-grained, moderately to well-sorted, cemented sandstones deposited mainly in braided ephemeral rivers and by eolian dunes. the skagerrak formation consists of interbedded sandstones, siltstones, claystones and anhydrites. the sandstones are arkosic, grey, red, orangebrown, fineto coarse-grained, poorly sorted, angular–subangular and partly cemented, and were deposited on alluvial fans or braided river plains. both formations are dominated by sandstones and potential internal barriers or seals are rare. another potential reservoir unit is represented by a muddy sandstone unit, 20–30 m thick, in the upper part of the lower jurassic f-ii member (fjerritslev formation) on the skagerrak–kattegat platform; the muddy sandstones were deposited by coastal progradation and ensuing transgression (sequences fj 4, fj 5; figs 8, 12; nielsen 2003). the sandstones are only well developed north-east of the børglum fault, where they show good porosity. in the south-eastern part of the study area, a series of lower jurassic (sinemurian–pliensbachian) shoreface sandstones were encountered in the lavø-1 and margrethe holm-1 wells, where they interfinger with marine mudstones of the fjerritslev formation. the sandstone unit is 30–70 m thick with porosities of 10–25%. the shoreface sandstones present in the lower and upper parts of the upper jurassic flyvbjerg formation are also considered potential reservoir rocks; these sandstones show a general thinning from north to south and from east to west (figs 8, 13). on the skagerrak–kattegat platform, the uppermost jurassic – lowermost cretaceous frederikshavn formation includes shallow marine and fluvial sandstones that possess reservoir properties, although the shale content increases rapidly towards the basin. 58 the evaluation presented here of potential source rocks in the danish portion of the norwegian–danish basin, together with a review of potential reservoirs, suggests that a mesozoic petroleum system may be present. two primary plays are possible: the upper triassic – lowermost jurassic gassum play and the middle jurassic haldager sand play, both relying on charge from lower jurassic (toarcian) or uppermost jurassic – lowermost cretaceous source rocks. both plays have, however, been tested with negative results in a number of wells. it is generally proposed that the main reason for the failure of these plays so far has been the insufficient maturation (burial depth) of the potential source rocks. in the light of this study, then, it is useful to revisit the important elements of the mesozoic petroleum system. source-rock quality and distribution the regional petroleum generation potential and thermal maturity of the pre-upper cretaceous succession in the study area have been assessed by evaluating the stratigraphic units drilled by 33 wells both onshore and offshore in the skagerrak and kattegat areas (fig. 1). it is generally accepted that the upper cretaceous – cenozoic strata have no sourcerock potential in the study area. within the lower palaeozoic – lower cretaceous succession, only the lower jurassic (toarcian) f-iii and f-iv members of the fjerritslev formation and the uppermost jurassic – lowermost creta ceous frederikshavn formation contain intervals that qualify as potential oil-prone source rocks in the successions drilled to date (figs 25–27, 34–38). none of these potential source rocks have a basinwide distribution. it is further emphasised that only parts of the lithostratigraphic units have a good to excellent petroleum generation potential and the potential source-rock units have highly variable generation potentials depending on the interaction of a number of geological processes during their formation. an overall upwards increasing petroleum generation potential is observed from the f-i member to the upper f-iii and f-iv members of the lower jurassic fjerritslev formation (figs 23, 24, 26, 27). this difference in generation potential is partly attributed to a change in depositional conditions through early jurassic times (thomsen et al. 1987; michelsen 1989b; nielsen 2003). the f-i and f-ii members were deposited in more oxic and shallow marine environments with a higher contribution of type iii kerogen, whereas the f-iii and f-iv members were deposited under more reducing bottom conditions with a higher contribution of oil-prone type ii kerogen (figs 28, 33). the generation potential of the toarcian part of the f-iii and f-iv members shows significant lateral changes, with the best-developed source-rock units occurring in the basin centre (fig. 35a; table 2). the source quality of the f-iii member seems to be particularly well developed in the mors-1, kvols-1, hyllebjerg-1 and farsø-1 wells, where the average hi values indicate a gas/oil generation potential (fig. 34a). the mors-1 well is located close to a major salt diapir (fig. 3), which may suggest that the good sourcerock quality is related to deposition in the deeper rim syncline of the diapir. it is likely that development of rim synclines adjacent to salt structures influenced source-rock formation, and it may thus be possible to infer the presence of source-rock intervals based on the analysis of lateral changes in seismic attributes in the rim synclines. indeed, thomsen et al. (1987) suggested that the increased hi values and toc contents in kvols-1 were the result of anoxic depositional conditions in a rim syncline; note, however, that this well was drilled adjacent to a minor salt pillow rather than a diapir. michelsen (1989b) proposed that the organic-rich section in the kvols-1 well resulted from reduced siliciclastic influx and constant organic deposition (i.e. a condensed section) as reflected in the relatively thin succession between ts 15 and the base of the f-iv member compared to other wells; this interpretation is compatible with the fact that the highest hi values seem to be related to mfs 15 (fig. 29a). the hyllebjerg-1 and farsø-1 wells were not drilled close to salt structures, but both wells are located in the deepest part of the himmerland graben (figs 2, 3). in hyllebjerg-1, the highest hi values actually occur in the basal part of the f-iv member immediately above sb 16, which may suggest a relationship to a transgressive surface (fig. 32a). the rønde-1 well is located on the eastern flank of the voldum structure, but the f-iii member is not developed as a good source rock in this well. in contrast, the well-developed source rocks occur in the f-iv member and they do not seem to be associated with either maximum flooding or transgressive surfaces (fig. 29b). it is commonly assumed that the formation of marine black mudstones with high values of hi is associated with initial flooding of lowstand systems creating sediment-starved environments with sufficient nutrients for high organic production (e.g. wignall & maynard 1993). during continued transgression, organic-rich black discussion 59 shales may be preserved. the formation of oil-prone shales is also considered to occur at the time of maximum marine flooding, and they are often best developed in the upper part of the transgressive systems tract close below the maximum flooding surface (bohacs 1993; robison & engel 1993; pasley et al. 1993). it is clear, however, from figs 29–32 that no simple relationship exists in this case between the sequence stratigraphic key surfaces and high hi values. the data from the wells analysed in this study thus indicate that the potential oil source rocks occur in different intervals within the toarcian part of the f-iii and f-iv members. these intervals may be locally developed, and 2–4 stacked intervals occur in some areas. the combined f-iii and f-iv members in the rønde-1, kvols-1 and haldager-1 wells, for example, possess oil-prone net source-rock intervals (i.e. hi >200 mg hc/g toc) with average hi values ranging from 294–404 mg hc/g toc, whereas the net source-rock interval in other wells is less oil-prone (fig. 35a). the large variation in net/gross ratio shows that sedimentation rate was not the sole controlling parameter (fig. 35a), but rather that optimum conditions required the right combination between sedimentation rate and organic productivity. local depressions in the basin may further have favoured preservation of the organic matter by promoting stratification of the bottom waters. a detailed analysis of outcrops and cores from the lower toarcian posidonia shale in south-west germany has shown that the formation of this rich source rock was governed by a complex interplay of factors, important amongst which was water column stratification controlled by sea-level changes (röhl et al. 2001). stratigraphic subdivision of the danish well sections (michelsen 1989a; nielsen 2003) indicates that the two lower intervals of organic-rich mudstones in the f-iii member are similar in age to the organic-rich shales in the falciferum and bifrons zones of the posidonia shale. this suggests that, in addition to local factors, external factors such as regional/global anoxic events and sea-level changes may have played an important role in the formation of organic-rich shales at this time in the norwegian–danish basin and the fennoscandian border zone. the regional or global anoxia in the early toarcian that seems to have favoured the formation of organic-rich mudstones at some stratigraphic levels in parts of the basin, was apparently not a significant factor in other areas, possibly owing to high clastic input or shallow water depth that prevented the development of oxygen-deficient conditions and thus masked the event. it is intriguing, however, that the anoxic event appears to have influenced deposition along the basin margin in shallow marine/lagoonal areas on bornholm (koppelhus & nielsen 1994; hesselbo et al. 2000). the implication is, therefore, that the deposition and preservation of marine organic matter in the basin was a complex interplay between a number of factors probably including bottom topography and depth, variation in primary organic productivity, sedimentation rate, bottom-water oxygenation and distance to fluvial sources and coastlines; the relative importance of these different factors in influencing source-rock formation in the area is poorly understood in detail. in contrast to the uncertainty related to regional prediction of source-rock quality, the present-day geographical occurrence of the stratigraphic interval spanning the f-iii and f-iv members is well understood (fig. 15). their occurrence is primarily dependent on the regional early middle jurassic uplift event that caused deep widespread erosion, with truncation of the source-rock interval over large parts of the study area. indeed, even within the central area where the f-iii and f-iv members are typically preserved (figs 15, 25b), this interval may be absent over local structural highs such as salt structures. to assess the distribution of this interval, it is necessary to map seismically both the base middle jurassic unconformity (base haldager sand formation or bryne formation) and the seismic reflector corresponding to top f-ii member. the latter coincides with a significant upper pliensbachian flooding surface typically situated 100–200 m below the best sourcerock intervals. if possible, reflectors between these two horizons should also be mapped out in order to further constrain the position of the potential source rocks. parts of the uppermost jurassic – lowermost cretaceous frederikshavn formation, which is dominated by siltstones and sandstones in most wells, possess a petroleum generation potential in the hyllebjerg-1, skagen-2, terne-1 and voldum-1 wells, with the terne-1 well having a particularly rich c. 160 m thick oil-prone interval with an average hi of 478 mg hc/g toc (fig. 38a; table 3). the type i kerogen composition of this terne-1 interval (fig. 39) indicates, however, freshwater to slightly brackish lacustrine depositional conditions in contrast to the marine and paralic conditions that characterised the regional depositional environment of the formation; the unit may thus be only a local development in the terne-1 area in the kattegat. source-rock maturity regional maturation profiles constructed from vr measurements corrected for post-early cretaceous exhumation yield a likely depth to the top of the oil window (vr of 0.6%ro) of c. 3050–3100 m, based on the regional coalification curves principally derived from onshore wells (figs 20a, 22a). accepting this depth, the potential source rocks 60 need a burial depth of c. 1.75–2 sec. twt to reach the oil window if velocity data from the års-1 well, placed centrally in the basin, are used as guidelines (nielsen & japsen 1991). the shales of the frederikshavn formation are thus regionally thermally immature in the study area. the toarcian shales of the f-iii and f-iv members of the fjerritslev formation constitute the most obvious potential source rocks, but over most of the study area they have not been buried sufficiently to have entered the oil window (fig. 40). of the investigated wells, only in the års-1 well, on the flank of the himmerland graben immediately west of a major salt structure, has the f-iii member been within the uppermost part of the oil window prior to post-early cretaceous exhumation. the source-rock quality of the f-iii member in this well, however, is poor with a maximum hi value of 186 mg hc/g toc and an average hi of only 113 mg hc/g toc. hence, occurrence of mature source rocks in the study area requires local burial anomalies, such as local grabens or rim synclines adjacent to salt diapirs, to reach thermal maturity for oil generation. a map showing the depth to the base of the middle juras sic (i.e. the base middle jurassic unconformity) in the danish part of the norwegian–danish basin was constructed by bidstrup et al. (2002). this surface corresponds to the top of the f-iv member, i.e. the top of the stratigraphic unit with the most promising source rocks (f-iii and f-iv members) in the fjerritslev formation. if the present-day depths are corrected for post-early cretaceous uplift (table 1; japsen 1998) and information of the combined thickness of the f-iii and f-iv members (fig. 25) is included, it is possible to locate potential areas with mature source rocks (fig. 41). these potential, locally developed petroleum kitchens are mainly located in the central part of the study area (central–northern jylland), where they are associated with rim synclines of salt structures. offshore, in the skagerrak, a minor kitchen may be present in the fjerritslev trough close to the fjerritslev fault. farther to the west, a kitchen may occur between two salt structures (fig. 41) although in this area the f-iii member is probably very thin and the f-iv member is absent (fig. 25). the onshore petroleum kitchens indicated on fig. 41 correspond to the possible kitchens in the harboør –uglev, mors and tostrup rim synclines mapped by thomsen et al. (1987); apart from the kitchen area in the mors rim syncline, however, the kitchens identified in the present study are areally smaller. the potential kitchen in the vejrum rim syncline suggested by thomsen et al. (1987) cannot be confirmed in the present study, principally due to very thin source-rock units (e.g. mejrup-1 well; fig. 25a) or the absence of the f-iii and f-iv members (fig. 15). 0.9 10.80.70.60.50.40.30.2 4000 3000 2000 1000 0 vitrinite reflectance (%ro) range of depths to the base of the f-iv member, based on 14 studied wells range of depths to the base of the f-iii member, based on 15 studied wells d ep th co rr ec te d (c ha lk ) ( m ) mors-1, base f-iii mb top of the oil window års-1, base f-iv mb års-1, base f-iii mb fig. 40. the regional vr curve based on chalk velocity corrected depth (see fig. 22a). the depth to the top of the oil window at 0.6%ro is indicated together with the range of depths to the base of the f-iii and f-iv members before postearly cretaceous exhumation in the studied wells. in the mors-1 well, the base of the f-iii member has been buried slightly deeper, but still above the top of the oil window. only in the års-1 well has the f-iii member been within the top part of the oil window before exhumation. the thickness of the f-iii member ranges from 30–279 m in the 15 studied wells, whereas the f-iv member ranges from 9–127 m in 13 of the 14 studied wells. in gassum-1, the f-iv member is considerably thicker, namely 320 m. 61 the mors-1 well is located adjacent to the northern flank of the largest potential kitchen area. the main targets of this well were the zechstein and rotliegend, but the well reached td in lower triassic sandstones at a depth of 5303 m. the well intersected 123 m of net source rock with gas and oil generation potential (average hi = 221 mg hc/g toc) and the source rocks are close to being early mature (fig. 35). the well encountered only small traces of asphalt in the haldager sand formation, and the poor hydrocarbon indications may be explained by the lack of structural closure (thomsen et al. 1987; geus, unpublished data). reservoirs and migration both reservoirs, the gassum and haldager sand formations, are proved to be present regionally and sealed by lower jurassic and upper jurassic mudstones, respectively. the haldager sand reservoir conformably overlies the lower jurassic potential source rocks in the relatively deep fjerritslev trough and presumably also in the farsund basin (fig. 2). in contrast, in the remaining parts of the study area, where the f-iii and f-iv members and the haldager sand formation are present, the haldager sand reservoir unconformably overlies the succession with the potential source rocks. migration of hydrocarbons to the haldager sand reservoir is thus simple. migration to the gassum reservoir requires stratigraphic downward migration of the hydrocarbons and may thus require structural components, i.e. faulting or salt domes. thin fine-grained shoreface sandstones formed during short-lived regressive events occur within the lower jurassic marine mudstones in the fjerritslev trough and skagerrak –kattegat platform. these sandstones can be traced relatively far into the basin as thin silty sandstone intercalations in the mudstones and may function as conduits for hydrocarbon migration. the gassum formation reservoir is overlain by the laterally consistent, thick marine mudstone succession of the fjerritslev formation. sandstone or siltstone units up to 5–10 m thick are present in the basal fjerritslev formation in places, but their influence on seal integrity over the gassum formation reservoir is generally expected to be limited. close to the basin margin, however, for example on the skagerrak–kattegat platform in the eastern part of kattegat and sjælland, hettangian–pliensbachian sandstones are common and constitute an additional potential reservoir, overlain by marine mudstones of the upper fjerritslev formation. the haldager sand reservoir is overlain by marine mudstones of the flyvbjerg and børglum for mations. sandstones are present in the lower and upper part of the flyvbjerg formation in places, and their thickness and grain size are expected to increase towards the northern and eastern basin margin, where they may form an additional reservoir section. the middle part of the flyvbjerg formation is dominated by marine mudstones with some seal capacity, and the flyvbjerg formation itself is overlain by the thick, regionally continuous, marine mudstone succession of the børglum formation. 100 km potential kitchen area zechstein salt diapir or pillow 57ºn 56ºn 6ºe 8ºe 10ºe 12ºe r i n g k ø b i n g – f y n h i g h r i n g k ø b i n g – f y n h i g h s o r g e n f r e i – to r n q u i s t z o n e s o r g e n f r e i – to r n q u i s t z o n e n o r w e g i a n – d a n i s h b a s i n n o r w e g i a n – d a n i s h b a s i n n o r w e g i a n – d a n i s h b a s i n r i n g k ø b i n g – f y n h i g h east north sea high farsund basin fjerritslev trough s o r g e n f r e i – to r n q u i s t z o n e hi. gr. fig. 41. map showing the location of potential petroleum kitchen areas with mature source rocks of the f-iii and f-iv members of the lower jurassic fjerritslev formation. the kitchens are locally developed and are principally associated with salt structures. the most westerly kitchen in the skagerrak is highly uncertain due to the absence of the f-iv member and possibly only a very thin (or absent?) f-iii member. note that only salt structures associated with potential kitchen areas are shown (cf. fig. 3). hi. gr., himmerland graben. 62 active mesozoic petroleum system? in agreement with the general immaturity of the potential source rocks in the study area, very few oil shows have been reported. one exception is the k-1 well (fig. 1), in which weak shows in sandstone and sandstone stringers were noted at several depths. the well was, however, drilled with diesel that was added several times during drilling operations; geochemical data of cuttings samples from the k-1 well show the presence of a low-boiling distillation cut such as diesel, with a minor contribution representing indigenous immature organic matter (geus, unpublished data). thus, reports of thermally generated petroleum in this well cannot be confirmed. the well was also drilled in an area in which the presence of mature source rocks is deemed unlikely (fig. 41). the presence of generated petroleum in the danish part of the norwegian–danish basin thus still has to be documented, although it cannot be excluded that petroleum has been generated in localised potential kitchen areas (fig. 41). conclusions two primary plays are possible in the study area: the upper triassic – lowermost jurassic gassum play and the middle jurassic haldager sand play, both relying on charge from lower jurassic (toarcian) or uppermost jurassic – lowermost cretaceous source rocks. both plays have, however, been tested with negative results. this study shows that two main uncertainties are present in the danish part of the norwegian–danish basin and the fennoscandian border zone: (1) the patchy distribution of well-developed, oilprone, potential source rocks, and (2) the thermal maturity of the potential source rocks. the latter factor is considered here to be the most significant uncertainty in proving the integrity of this mesozoic petroleum system. the evaluation of source-rock quality, thermal maturity and distribution allows the following principal conclusions to be drawn: 1. lower palaeozoic rocks are overmature in the study area and upper cretaceous – cenozoic strata possess no petroleum generation potential. 2. toarcian marine shales of the lower jurassic f-iii and f-iv members of the fjerritslev formation and the uppermost jurassic – lowermost cretaceous shales of the frederikshavn formation constitute oil-prone potential source rocks in parts of the basin. the generation potential of these potential source rocks is highly variable geographically, and the f-iii and f-iv members in the centre of the basin possess the best-developed source potential. the highly oil-prone lacustrine mudstone interval of the frederikshavn formation in the terne-1 well is probably only a local development. 3. based on interpretation of regional coalification curves, the top of the oil window (vitrinite reflec tance = 0.6%ro) is located at c. 3050–3100 m depth. the uppermost jurassic – lowermost cretaceous frederikshavn formation had not been buried to this depth prior to post-early cretaceous exhumation, and the potential source rocks of the formation are thermally immature in terms of hydrocarbon generation. similarly, the potential source rocks of the lower jurassic f-iii and f-iv members of the fjerritslev formation are generally immature to very early mature. however, potential kitchen areas with mature source rocks of the f-iii and f-iv members may occur in the central part of the study area (central–northern jylland) and a few places offshore. these potential petroleum kitchens are considered to be of local development, mainly associated with salt structures and grabens (fjerritslev trough and himmerland graben). acknowledgements the study was in part supported financially by the danish energy authority and dong energy. a.m. spencer and an anonymous referee are thanked for constructive reviews which greatly improved the manuscript. 63 andsbjerg, j., nielsen, l.h., johannessen, p. & dybkjær, k. 2001: divergent development of depositional environments in the danish central graben and the norwegian–danish basin following the jurassic north sea doming event. in: martinsen, o.j. & dreyer, t. 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(eds): petroleum geology of north west europe, 497–508. london: graham & trotman. vejbæk, o.v. 1989: effects of asthenospheric heat flow in basin modelling exemplified with the danish basin. earth and plane tary science letters 95, 97–114. vejbæk, o.v. 1997: dybe strukturer i danske sedimentære bassiner. geologisk tidsskrift 4, 1–31. vejbæk, o.v. & britze, p. 1994: geological map of denmark, 1:750 000. top pre-zechstein. danmarks geologiske undersøg else kortserie 45, 6 pp. (unnumbered). wignall, p.b. & maynard, j.r. 1993: the sequence stratigraphy of transgressive black shales. in: katz, b.j. & pratt, l.m. (eds): source rocks in a sequence stratigraphic framework. american asso ciation of petroleum geologists studies in geology 37, 35–47. zdanaviciuté, o. & bojesen-koefoed, j.a. 1997: geochemistry of lithuanian oils and source rocks: a preliminary assessment. journal of petroleum geology 20, 381–402. ziegler, p.a. 1982: geological atlas of western and central europe, 130 pp. the hague: shell inter nationale petroleum maat schappij b.v. ziegler, p.a. 1990: geological atlas of western and central europe, 2nd edition, 239 pp. the hague: shell internationale petroleum maatschappij b.v. geological survey of denmark and greenland bulletin 15, 2008, 9-72 69 tonalite-trondhjemite–granodiorite (ttg) gneisses and mela nocratic to ultramafic greenstones dominate the ar chaean basement of southern west greenland. the greenstones are likely to represent different original environments, which is important as the mineral deposits they may host depend on this. for example, massive sulphide deposits associated with gold and base metals are commonly volca no genic, while chrome, nickel and platinum group elements are more commonly associated with layered intrusions (robb 2005). cur rent investigations by the geological survey of denmark and greenland (geus) in southern west green land are therefore focused on the origin of greenstones and their relationship to associated ttg gneisses. here, we report on work in progress on greenstones within the tasiusarsuaq terrane (fig. 1; friend et al. 1996). they differ from many other greenstone belts in southern west green land in their spatial association with the ttg gneisses. unlike the isua, ivisârtoq and storø greenstone belts in the central and northern nuuk region, the tasiusarsuaq greenstones are not proximal to terrane boundaries but form dismembered blocks and slivers within the terrane (fig. 1). contact relationships to the gneisses are almost exclusively tectonic, and primary textures are, with rare exceptions, ob literated by amphibolite to granulite facies metamorphism. field relationships stendal & scherstén (2007) documented one of the rare examples of well-preserved field relationships for a volcanic pile of pillow basalts, rhyolites (sensu lato) and melanocraticultramafic tuffs and flows on ‘nunatak 1390’ (fig. 1). the rhyolite was extruded at 2.876 ± 0.005 ga (næraa & scherstén 2008 – this volume), which is the minimum age for the basalts and within the known age range for the tasiusarsuaq terrane. geochemistry of greenstones in the tasiusarsuaq terrane, southern west greenland anders scherstén, henrik stendal and tomas næraa © geus, 2008. geological survey of denmark and greenland bulletin 15, 69–72. available at: www.geus.dk/publications/bull greenland tre brødre færingehavn terrane tasiusarsuaq terrane terrane ‘nunatak 1390’ nuuk serm ilik 50°52° 64° 25 km qarliit nunaat thrust bu kse fjo rde n go dt hå bs fjo rd ? ameralik akia terrane kapisilik terrane nunataarsuq meso–neoarchaean granites meso–neoarchaean orthogneiss and granitic rocks. dots: granulite facies supracrustal belts (undifferentiated) anorthosite-gabbro complexes qôrqut granite complex terrane boundary palaeoproterozoic fault structural trend line eoarchaean orthogneiss study areas fig. 1. generalised map of the northern part of the tasiusarsuaq terrane. the northern boundary as suggested by friend et al. (1996) is outlined with solid and stippled lines for more or less well-determined demarcations of the terrane. it is likely that the border extends farther to the north in the eastern part (cf. næraa & scherstén 2008 – this volume; b. windley and a.a. garde, personal communication 2008). sampling areas are highlighted by red boxes, and the two main sampling areas are indicated by solid lines. 70 here, we make the assumption that the rhyolite date establishes the age for the majority of the greenstones within the tasiusarsuaq terrane. to the north-west of ‘nunatak 1390’, a melanocratic to ultramafic complex forms a tectonic lens, with internal topto-the-north thrust planes that reflect tectonic emplacement against the gneisses (figs 1, 2a). the rocks range from ultramafic olivine-pyroxene-rich cumulates to basaltic amphibolites, sometimes with flattened pillow lavas and calc-silicates (fig. 2b). ultramafic rocks with pillow-like structures were also observed, and the complex is interpreted to be dominated by basaltic to komatiitic flows and shallow sills (fig. 2c). similar dismembered bodies are scattered within the tasiusarsuaq terrane but are commonly more deformed and do not preserve primary textures. the ultramafic rocks have been divided into two groups on the basis of their field appearance, with the first group containing ultramafic rocks of cumulate or undetermined origin and the second comprising ultramafic rocks with eruptive features such as pillow structures. only the eruptive ultramafic rocks will be considered here. whole-rock geochemistry the rocks considered here are dominated by tholeiitic metavolcanic amphibolites and meta-ultramafic rocks of ko matiitic composition. the ultramafic rocks are signified by mgo >16 wt%, al2o3 <5.5 wt%, high ni and cr (>500 and >1500 ppm, respectively), and al2o3/tio2 ratios <10, while the amphibolites have mgo <10 wt%, al2o3 >10 wt%, ni and cr (<250 and <500 ppm, respectively) and al2o3/tio2 ratios of 7–37, of which most ratios are >10 (fig. 3). the ultramafic rocks are aluminium depleted and show strong positive correlation between ni and mgo (fig. 3), which is presumably controlled by olivine and pyroxene fractionation. al2o3/tio2 ratios of 10–40 for the amphibolites and <10 for the ultramafic eruptives imply a deeper melt origin (>5 gpa, i.e. >150 km) for the ultramafic rocks (walter 1998). on ‘nunatak 1390’ at least some of the ultramafic rocks seem to be intercalated with basaltic pillow lavas, which might suggest that they formed sills that are younger than the basalts (stendal & scherstén 2007). this could imply that the depth of melting increased with time (presuming that the ultramafic rocks are indeed slightly younger than the pillow lavas), or that shallow melting was induced by the deeper melts if they were contemporaneous. the ultramafic rocks have smooth trace element patterns, but with nb/th ratios that are lower than the primitive mantle (pm; fig. 4). this implies some degree of enrichment in the mantle source, or that the nb/th ratios have decreased due to shallow continental crustal contamination. the light rare-earth element (ree) signatures are horizontal, whereas the midto heavy ree signature slopes towards lower abundances, indicating a garnet residue during mantle melting, which is consistent with their deep origin as discussed above (fig. 4). the ree signatures of the amphibolites are similar to those of the ultramafic rocks, but with slightly higher conn a b c fig. 2. a: tasiusarsuaq greenstones displaying a tectonic contact with the tasiusarsuaq grey ttg gneises; the greenstones seem to have been thrust northwards. b: variably deformed pillows, sometimes with elongate calc-silicate aggregates, demonstrating a supracrustal origin for the greenstones. such rocks are found throughout the area (cf. stendal & scherstén 2007, fig. 4). the flattened pillows are cross-cut by a granite dyke of unknown age. c: pillow-like structures among ultramafic rocks. centrations, and lacking the garnet signature noted for the ultramafic rocks (fig. 4). nb/th ratios are generally pm-like. overall, the amphibolites are generally more mid-oceanic ridge basalt (morb)-like in their range of trace element ratios and abundances. in particular, one core of a pillow basalt from ‘nunatak 1390’ consistently lacks continental crustal (or arc-like) trace element patterns. crustal contamination the general geochemical signatures and the subaqueous nature of the amphibolites and ultramafic rocks are consistent with an ocean-floor origin in its broadest sense. the morb-like signatures of the amphibolites might be suggestive of an ocean basin, a back-arc basin or even a primitive tholeiitic arc. however, a major extensional setting such as a mid-ocean ridge is at odds with major simultaneous continental crust formation, which is indicated by e.g. massive volumes of presumably contemporaneous ttg crystallisation (næraa & scherstén 2008 – this volume), while a backarc basin seems more conceivable. the trace-element arrays indicate source enrichment, i.e. variable amounts of enriched subduction components, or local contamination during emplacement (fig. 4). positive correlations for nb/th and nb/la against the th or la concentration reciprocals lie between mantle and continental crustal end-members, and these ratios are the most sensitive to small degrees of conta71 15 15 5 1200 800 400 0 25 35 1500 500 1000 amphibolites ultramafic rocks 10 5 0 5 10 15 20 25 0 5 10 15 20 25 a l 2o 3 a l 2o 3/ t io 2 n ic r mgo mgo fig. 3. mgo bivariate plots for a selected oxide, an oxide ratio and compatible trace elements. see text for further discussion. primitive mantle ocean-island basalts (oib) th nb la ce nd sm zr eu ti gd tb dy er y yb 100 10 1 0 1.4 1.2 1.0 0.8 0.6 0.4 5 10 15 nb/th la/smn tonalite-trondhjemite-granodiorite (ttg) normal mid-ocean ridge basalts (n-morb) tonga arc basalts lau basin back-arc basalts amphibolites ultramafic rocks n-morb 1% mixing increments between a morb-like end member and archaean ttg pillow lava mean amphibolites mean ultramafic rocks fig. 4. primitive mantle (palme & o’neill 2004) normalised trace element diagram for tasiusarsuaq greenstones, one pillow lava from ‘nunatak 1390’ and selected reference rocks and reservoirs. for comparison plots of n-morb (hofmann 1988), oib (sun & mcdonough 1989), archaean ttg (martin 1995; martin et al. 2005), lau basin back-arc (regelous et al. 2008 ) and median tonga arc basalts (http://georoc.mpch-mainsz.gwdg.de /georoc) are shown. inset: n-morb, lau basin back-arc basalts and oib are char acterised by nb/th ratios that are higher than pm, while median tonga arc basalts and ttg have ratios that are lower than pm. the tasiusarsuaq ultramafic rocks and one pillow lava have arc-like ratios lower than pm nb/th, while the amphibolites are variable with both suband supra-pm ratios. minor ttg contamination of mag mas with morb-like ratios would rapidly decrease nb/th with associated increasing la/sm, as these ratios are extreme in ttg. a plot of nb/th against chondrite normal ised la/sm is shown in the figure inset for the tasiusarsuaq data, displaying a mode rate fit with a mixing scenario as discussed above. mination. assuming a ttg crustal component as the contaminant, the array can be explained by <5% contamination for all but one sample, supposing that the most primitive ul tramafic rocks are uncontaminated. tectonic implications if the basalt–komatiite magmatism in the tasiusarsuaq terrane is indeed concurrent with ttg-formation, an arc environment for the former magmatism is favoured (cf. stendal & scherstén 2007; næraa & scherstén 2008 – this volume), and such a hypothesis is still viable in the light of the current geochemical data. the origin of komatiites remains controversial, although most authors advocate a mantle-plume related origin. the komatiite-like rocks documented here do not readily fit such an origin as they seem to be primarily associated with subduction and growth of continent crust. alternatively, renewed models for subduction-related komatiite genesis might be considered. however, this scenario typically involves shallow melting (grove & parman 2004), while the ree ratios observed here favour deep melting with residual garnet. outlook further work with detailed field studies and geochronology over the next few years will hopefully shed new light on these outstanding issues. emphasis will be placed on searching for primary relationships between the ttg gneisses and the ultramafic rocks and amphibolites in conjunction with detailed geochronology and geochemistry. references friend, c.r.l., nutman, a.p., baadsgaard, h., kinny, p.d. & mcgregor, v.r. 1996: timing of late archaean terrane assembly, crustal thickening and granite emplacement in the nuuk region, southern west greenland. earth and planetary science letters 142, 353–365. grove, t.l. & parman, s.w. 2004: thermal evolution of the earth as recorded by komatiites. earth and planetary science letters 219, 173–187. hofmann, a.w. 1988: chemical differentiation of the earth: the relationship between mantle, continental crust, and oceanic crust. earth and planetary science letters 90, 297–314. martin, h. 1995: archean grey gneisses and the genesis of continental crust. in: condie, k.c. (ed.): archean crustal evolution, 205–260. amsterdam: elsevier. martin, h., smithies, r.h., rapp, r., moyen, j.-f. & champion, d. 2005: an overview of adakite, tonalite-trondhjemite-granodiorite (ttg), and sanukitoid: relationships and some implications for crustal evolution. lithos 79, 1–24. næraa, t. & scherstén, a. 2008: zircon age perspective on the ta siusarsuaq terrane, southern west greenland. geological survey of denmark and greenland bulletin 15, 73–76. palme, h. & o’neill, h.s.c. 2004: cosmochemical estimates of mantle composition. in: carlson, r.w. (ed.): treatise on geochemistry 2, 1–38. amsterdam: elsevier. regelous, m., turner, s., falloon, t.j., taylor, p., gamble, j. & green, t. 2008: mantle dynamics and mantle melting beneath niuafo’ou island and the northern lau back-arc basin. contributions to mineralogy and petrology. doi: 10.1007/s00410-007-0276-7. robb, l. 2005: introduction to ore-forming processes, 373 pp. oxford: blackwell science ltd. stendal, h. & scherstén, a. 2007: a well-preserved bimodal volcanic succession in the tasiusarsuaq terrane, south-west greenland. geological survey of denmark and greenland bulletin 13, 53–56. sun, s.-s. & mcdonough, w.f. 1989: chemical and isotopic systematics of oceanic basalts: implications for mantle composition and processes. in: saunder, a.d. & norry, m.j. (eds): magmatism in the ocean basins. geological society special publication (london) 42, 313–345. walter, m.j. 1998: melting of garnet peridotite and the origin of komatiite and depleted lithosphere. journal of petrology 39, 29–60. 72 authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: asch@geus.dk kortbladsbeskrivelse, geologisk kort over danmark map description pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 1 of 53 kortbladsbeskrivelse, geologisk kort over danmark, 1:50 000, møn dele af 1511 i, 1511 iv og 1512 ii with a summary in english stig a. schack pedersen*1 , peter gravesen1 1geological survey of denmark and greenland (geus), copenhagen, denmark. sammendrag det geologiske kortblad møn omfatter møn med de tilgrænsende øer langø, lindholm og nyord samt mindre dele af sjælland og falster. kortet består af dele af de topografiske kortblade 1511 i og 1512 ii samt 1511 iv med rand­ områder af tilgrænsende kortblade mod vest og nord. møn opdeles i tre geomorfologiske områder: det stærkt kuperede høje møn mod øst, det småbakkede landskab omkring stege nor mod vest, og det flade marine forland omkring nyord og ulfshale. høje møn opbygges af opskudte skiver af skrivekridt og kvartære aflejringer, som det ses i møns klint. skiverne er op til 80 m tykke, hvoraf skrivekridtet udgør ca. 50 m. under hele møn består prækvartæroverfladen af maastrichtien skrivekridt i en dybde omkring kote –25 til –40 m. mindre skiver af glacialtektonisk forstyrret skrivekridt optræder også omkring stege nor og langs sydkysten af det vest­ lige møn ved hvideklint. de ældste kvartære aflejringer er moræneler fra saale­istiden og sand og ler fra eem­mellemistiden. derefter følger fluviale aflejringer og nedskylslag fra tidlig weichsel. disse lag efterfølges af moræneler fra ristinge klint till formationen med over­ og underliggende smeltevandsaflejringer fra mellem weichsel dannet under ristinge isfremstødet for ca. 55 000–50 000 år siden. den næste enhed er kraneled formationen (ny formation), som efterfølges af moræneler tilhørende klintholm till formationen (justeret formation) fra klintholm isfremstødet for 35 000–32 000 år siden. formationen overlejres af mere end 10 m tykke enheder af gråt til olivengråt issøler med dropsten, smeltevandssand og lamineret fint sand samt diamikte aflejringer i kobbel­ gård formationen (ny formation). denne formation blev aflejret i en issø, som dækkede store dele af østersøen i en mildningsperiode for 32  000– 28  000 år siden. denne enhed overlejres af eller er øverst sammenflettet med sand og grus tilhørende stubberup have formationen (ny formation). moræneler tilhørende den midtdanske till formation blev aflejret under nø­isfremstødet for 23 000–20 000 år siden. efter at nø­isen var smeltet til­ bage fra østersøområdet, rykkede den ungbaltiske is frem fra den østlige del af østersøen, hvorunder bl.a. møns klint og hvideklint blev deformeret. en tilhørende strukturel enhed, møns klint glacialdynamiske kompleks, er defi­ neret med fire sekvenser. hele hjelm bugt dannede en glacial lobe, og nord herfor dannedes et randmorænestrøg. radialt ud fra loben dannede smelte­ vandet store afløbskanaler fra gletsjerporte i den ungbaltiske is. aflejring af *correspondence: sasp@geus.dk received: 04 feb 2019 accepted: 18 apr 2021 published: 14 dec 2021 keywords: geologisk kort, kvartær geologi, glacialtektonik, skrivekridt, fjeldskred geological map, quaternary stratigraphy, cretaceous chalk, glacial tectonics, landslides abbreviations: geus: geologiske undersøgelser for danmark og grønland hk: hvideklint irsl: infrared stimulated luminescence mis: marine isotope stage osl: optisk stimuleret luminiscens lgm: seneste glaciale maksimum tl: thermoluminescence geus bulletin is an open access, peer­reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: adam a garde (geus, denmark) reviewed by: helena alexanderson (lund university, sweden) & nikolaj krogh larsen (university of copenhagen, denmark) funding: see page 50 competing interests: none declared additional files: see page 50 pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 2 of 53 www.geusbul let in.org sand og grus tilhørende ny borre formationen (ny formation) skete i dette tidsrum. under det ungbaltiske isfrem­ stød blev lolland till formationen aflejret som et relativt tyndt lag af moræneler. ved slutningen af weichsel­istiden for ca. 17  000 år siden smeltede den ungbaltiske is tilbage. et residualt isdække i området nordøst for møn sendte et genfremstød til det østlige møn, som medførte en reorientering af skrivekridtskiverne i møns klint. i sen weichsel (17 000–11 700 år før nu) fandtes søbassiner på det sydlige møn ved hjelm og tøvelde samt på høje møn, hvor en række ferskvandslag blev dannet, og aflejringen fortsatte et stykke ind i holocæn. i holocæn blev de tidligere afløbskanaler transgrederet under den atlantiske havstigning, hvorved fjorde skar sig ind fra nord og nordvest til midt på møn. herefter begyndte udbygningen af marine forlande, især mod nord i områ­ det ulvshale og nyord. de tidligere fjorde voksede til med planter, som omdannedes til tørveaflejringer. den sidste sedimentationsfase skete langs kysterne, hvor strandvolde blev akkumuleret, og kystklitter af flyvesand blev dannet. 10°ø 56°n kattegat skagerrak femer bælt østersøen storebæ lt lolland la ng el an d falster møn ristinge klint fyn jyl land sjæl land sverige danmark 50 km ø resund fig. 1 indekskort over danmark med be ­ liggenheden af kortbladet møn. fig. 1 index map of denmark with the position of the map sheet møn. introduktion topografi det geologiske kortblad møn i skala 1:50  000 omfatter den nordlige halvdel af det topografiske kort 1511  i og den sydlige halvdel af 1512  ii samt en mindre del af 1511  iv (fig. 1). kortbladet dækker øen møn, farvandet omkring møn med øerne nyord, lindholm og langø, samt mindre dele af sydøstsjælland og nordøstfalster. området er beliggende mellem 11°15’–11°42’  ø og 54°47’–55°04’  n. alle stednavne nævnt i teksten findes på det geologiske kort eller figurerne. kortet er indsat bagerst i kortbladsbeskrivelsen. kortbladsdelen på sydøstsjælland omfatter et mindre trekantet område, som strækker sig fra kalvehave mod nord til sandvig kohave med viemose skov og fra kalve­ have mod vest til langebæk. mod vest når området op til 46 m o.h. og består af en morænelerflade hældende mod kysterne, hvor holocæne havaflejringer af ler og http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 3 of 53 www.geusbul let in.org sand findes flere steder, især ved kindvig hoved. stege bugt med øen lindholm ligger mellem sjælland og møn, mens den større langø ligger nord for ulvsund. på nordøstfalster er et lille område ved grønsund nord for korselitse østerskov og øst for stubbekøbing en del af kortområdet. det består af moræneler og inddæmmede holocæne leraflejringer i noret. ud mod grønsund færgebro og kysten sydpå mod grønsund består vinkelforlandet også af havaflejringer, især sand, grus og rullesten af flint (klint et al. 2017). ud fra højdeforholdene kan møn inddeles i tre karak­ teristiske terræntyper: 1) høje møn, 2) det glacialmor­ fologiske landskab omkring stege nor, og 3) det flade marine forland omkring nyord og ulvshale. det højeste punkt på møn ligger i store klinteskov, som er det skovområde, der karakteriserer høje møn. højdepunk­ tet med topkote på 143 m o.h. ligger på aborrebjerg, der er et markant bakkedrag oven for aborresø. aborre bjerg er dog ikke det eneste højdepunkt i store klinteskov. adskillige toppunkter (kongsbjerg, lollikebakke, linde­ bakker og store ørnebjerg) når højder over 130 m, og her findes danmarks højest beliggende søer aborresø, store geddesø og hunesø med vandspejle tæt på 90 m o.h. det mest dramatiske højdepunkt er dog dronninge­ stolen, som med sin højde på omtrent 130 m har et næsten lodret fald direkte til havet. den n–s­gående lavning ved borre sømose og buse­ marke mose, der stedvis ligger i kote 1 m o.h., adskiller høje møn fra det glacialmorfologiske landskab, der omkranser stege nor. imellem elmelunde og sønderby/borre har bakkelandet stadig en vis højde, men kommer dog aldrig over kote 50 m. det glacialmorfologiske landskab ligger generelt omkring kote 25 m, men hist og her optræder bakkekomplekser med højder op til kote 40 m. bortset fra store områder med holocæne aflejringer: kostervig, nyord, ulvshale, maglemose, borre sømose, busemarke mose, råbylille sø og råby sø, er møns overflade domineret af glaciale weichsel­sedimenter med stort indhold af skrivekridt. især høje møn – møns klint og hvideklint er eksponenter for glacialtektoniske klinter med store skrivekridtskiver mellem moræne­ og smeltevandsaflejringer. særligt på høje møn kan dette registreres på terrænoverfladen. det flade marine forland ved ulvshale ligger tæt ved havniveau med højder på mellem 1 og 3 m o.h. de højeste områder med topkote på 3 m er langstrakte bakke rygge med flyvesand aflejret på strandvolde. til disse områder hører tillige det inddæmmede område ved klostervig samt gamle fjordområder som borre sømose og maglemose. endelig er møns sydpynt ved hårbølle havn opstået som et marint vinkelforland. store dele af kysterne består af udlignede kystklinter afbrudt af marint forland, især syd­ og østkysterne, men også den østlige del af nordkysten. mod vest fra ulvshale og sydpå er vestkysten guirlandeformet med stege nor og fanefjord som de største indbugtninger. hjelm bugt, fakse bugt og stege bugt, der omgrænser møn, er fladvandede havområder, som skæres af få dybe sunde som grønsund (vanddybde op til 23 m) mellem møn og falster og ulvsund (vanddybde op til 18 m) mellem sjælland og møn. der findes kun få mindre øer inden for disse havområder, hvoraf nyord og langø er de største. metode den systematiske geologiske kortlægning af den dan­ ske landoverflade betragtes traditionelt også som en kortlægning af de kvartærgeologiske forhold, da mere end 90 % af danmarks landoverflade består af kvartære aflejringer. kortlægningen udføres af de nationale geo­ logiske undersøgelser for danmark og grønland (geus) og foregår ved, at geologer udstyret med karteringsspyd gennemvandrer landskabet. for hver ca. 100 m fore­ tages en sondering med et karteringsspyd ned til ca. 1 m dybde, således at den kortlagte jordart repræsen­ terer lagene under eventuelle kulturlag og jordbunds­ horisonten. prøverne, som opsamles i en lille slids i jernspyddet, bestemmes på stedet, og jordartstypen indføres på karteringskortet. princippet i kortlæg ningen er dels at kortlægge jordartsgrænser i felten, dels at sikre en tilstrækkelig dækning af sonderinger inden for hver kortlagt jordartspolygon. ofte foretages sonderin­ gerne med tættere afstand, især hvor de geologiske grænser skal fastlægges (gravesen et al. 2006; jakobsen et al. 2015). informationer fra vejskæringer, råstofgrave, andre udgravninger og kystprofiler inddrages også i kortlægningen. i felten skelnes mellem ca. 40 forskellige jordarts­ typer. disse har hver et symbol, som indtegnes på felt­ kort i skala 1:25 000 (jakobsen et al. 2015; se appendiks i pedersen et al. 2015). disse foreløbige jordartskort danner grundlaget for det publicerede geologiske kort. de digitaliserede foreløbige jordartskort findes i geus’ jordartsdatabase og kan hentes på geus’ hjemmeside (www.geus.dk). grundlaget for etablering af litostratigrafiske en heder i kvartære aflejringer er beskrevet i pedersen et al. (2015). desuden er der ved de kvartære enheder angi­ vet, hvilket ’marine isotope stage’ (mis) aflejringerne relateres til (cohen & gibbard 2012). alle aldersangiv­ elser i denne kortbladsbeskrivelse følger houmark­niel­ sen et al. (2017). dataindsamlingen fra den systematiske geolo­ giske overfladekartering opbevares i geus’ arkiver. langt den over vejende del af landarealet vist på kort­ bladet er dækket af moræneler, i alt 70 %, som på selve møn udgør 218 km2 foruden mindre arealer på sydsjælland og østfalster. de sandede gla­ ciale aflejringer domineres af smeltevandssand og http://www.geusbulletin.org http://www.geus.dk pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 4 of 53 www.geusbul let in.org ­grus, som træder frem i overfladen på 11 % af land­ arealet. de resterende dele dækkes af holocæne fersk­ vandsaflejringer (7 %), holocæne marine aflejringer (10 %) samt skrivekridt (1 %), især på høje møn. byer og befæstede anlæg (som ikke er karteret) samt fersk­ vandssøer udgør 1 %. siden ca. 1985 er der udført en hel del videnskabe­ lige undersøgelser af klinterne på syd­, øst­ og nord­ møn, herunder flere specialestudier. desuden er der foretaget undersøgelser af råstof­ og grundvandsfore­ komster, som er offentliggjort i rapporter. resultater fra alle disse undersøgelser er i vid udstrækning inkluderet i denne kortbladsbeskrivelse. kortlægningens historie de første geologiske observationer på møn blev offentlig­ gjort af pontoppidan (1768), som med et træsnit demon­ strerede den geologiske opbygning af den sydligste del af møns klint fra vidskud over græderen og sommer­ spiret til dronningestolen (se afsnit om møns klint). puggaard (1851, 1852) foretog en omhyggelig opmåling af hele møns klint og beskrev herudfra den geologiske opbygning af møn. selv om puggaard ensidigt støttede en istektonisk dannelsesmekanisme for møns klint, blev det først med johnstrups (1874) afhandling almindeligt accepteret, at flageopskydningen hidrørte fra et ispres forårsaget af en baltisk isstrøm, der ved sin ø–v­rettede strøm gennem østersøen både påvirkede rügen og møn, hvorved skrivekridtflagerne med deres morænedække blev skubbet op. en kortlægning af møn blev allerede forestået af k.a. grønwall i 1906. hintze (1937) udfærdigede et kort over høje møn, som fastholdt en blok­tektonisk opfattelse af deformationerne i møns klint. samtlige nyere geolo­ giske arbejder understøtter en glacialtektonisk årsag til deformationerne (konradi 1973; berthelsen et al. 1977; pedersen 1988, 2000, 2011; houmark­nielsen 1994). på sit kort har hintze (1937) tolket samtlige afløbsløse huller som jordfaldshuller, altså huller opstået som en slags karstfænomener, men langt de fleste er dødishuller, spor af begravede islegemer efterladt i landskabet, da isdækket fra den ungbaltiske is smeltede bort. milthers (1948) sammenstillede et kvartærgeologisk kort over hele danmark i målestok 1:320 000 og inddrog også møn. på møn er der dog kun differentieret mel­ lem morænelandskab, inddæmmede havaflejringer og strandvoldskomplekser. v. haarsted præsenterede ved et møde i dansk geologisk forening i januar 1955 et geomorfologisk ter­ rænkort og et jordbundskort i målestok 1:20 000 over møn. en sammenfatning af de geologiske forhold findes i et oversigtskort trykt i referatet af foredraget (haarsted 1956), hvor den glacialgeologiske dannelse blev under­ streget af bakkestrøg vist og tolket som ispreslinjer. desuden er dalsænkninger og åsdannelser markeret. i en kompilation af jordartskort over danmark (pe der­ sen 1989) i målestok 1:200 000 blev der også inddraget oplysninger fra boringer og flyfotos. herved blev den østligste del af høje møn udlagt som et område med skrivekridt, og et mindre antal skrivekridtflager blev angivet på det østligste møn. den kortlægning, som er grundlaget for det nær­ værende kortblad over møn, blev udført i flere perioder af forskellige geologer. v. haarsted, der var rektor på holte gymnasium, afsluttede sin kortlægningsindsats i 1959. v. münther og a.v. nielsen kortlagde mindre arealer på møn i årene 1966–1972. p. gravesen og s.a.s. pedersen afsluttede kortlægningen i årene 2004–2006, men havde forinden udarbejdet et foreløbigt kort over det østlige møn (pedersen & gravesen 2005). -20 -20 -20 -20 -10 -10 -10 -10 -30 30 -30 -30 10 10 50 80 0 0 0 0 0 n 5 m fig. 2 topppen af skrivekridtet under de kvartære lag, herunder dislocerede skrivekridtflager. højdekurver vist med 10 m intervaller over og under havniveau. fig. 2 the top of chalk below the quaternary deposits, including chalk sheets dislocated by glacial tectonics. contour lines with 10 m intervals shown above and below sea level. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 5 of 53 www.geusbul let in.org prækvartære aflejringer under de pleistocæne aflejringer på møn findes skrivekridt fra sen kridt (for 99,6–66,0 mio. år siden). møn er beliggende på den sydlige del af ringkøbing– fyn højderyggen (sorgenfrei & buch 1964; sorgenfrei 1966: håkansson & pedersen 1992; surlyk et al. 2013). denne beliggenhed gør, at der blev aflejret ca. 500 m tykke kalklag omkring og under møn, mens der nordfor i det danske bassin er op til 2000 m tykke kalkaflejringer (surlyk et al. 2013). størstedelen af disse kalkaflejringer er skrivekridt. syd for højderyggen er der også fundet forholdsvis tynde (<500 m) lag af skrivekridt i det tyske bassin (surlyk et al. 2013). der er ingen dybe undergrundsboringer på møn. den nærmeste dybe boring findes på falster, hvor en 2,5 km dyb boring, ørslev­1, blev gennemført ved ørslev af olieselskabet gulf i 1967–1968 (nielsen & japsen 1991; geus 1997). i den dybeste del af boringen blev der truf­ fet lag fra karbon, og ud fra seismiske data må det for­ modes, at grundfjeldet ligger i en dybde på mindre end 3 km under terrænoverfladen. lag af stensalt, som på lolland er trængt op i puder og mindre diapirer, findes i en dybde af 1,5 km. i ørslev­1­boringen findes bunden af skrivekridtet i en dybde af 450 m. skrivekridt fra sen kridt skrivekridtet er tilgængeligt sammen med og indlejret i de kvartære aflejringer mange steder langs møns kyst i a b fig. 3 skrivekridt med sorte flintlag. a: stærkt foldede flintlag i st. stejlbjerg klinten, sydlige møns klint. klinten er her knap 100 m høj. b: nærbillede af sorte flintlag i skrivekridt. den ring­ formede flintkonkretion nederst til venstre tolkes som udfældet omkring en paramoudra gravegang i kridthavets bund. fig. 3 pale grey chalk with black flint layers. a: folded flint layers in the chalk cliff of st. stejlbjerg, southern møns klint. the chalk cliff is here almost 100 m high. b: close-up of black flint layers in chalk. the large, circular flint concretion in the lower left of the photo is interpreted as formed around a paramoudra burrow in the cretaceous sea floor. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 6 of 53 www.geusbul let in.org de glacialtektonisk påvirkede kystprofiler, men træffes kun faststående i boringer. boredata viser, at præ­ kvartæroverfladen på store dele af møn ligger mellem 25–40 m u.h. og udgør en relativt plan flade, som kun stedvis når ned til større dybder (se profiler på kort­ bladet samt i houmark­nielsen 2003). beliggenheden af overfladen under de opskudte skiver på høje møn er dårligt kendt, men er muligvis dybere end på det øvrige møn. i området nord for hvideklint når overfladen helt op omkring kote 0 m, men det skyldes sandsynligvis en del opskudte flager. fra området omkring elmelunde til borre ligger prækvartæroverfladen også i kote 0 m, for­ modentlig af samme grund (fig. 2). skrivekridtet er en hvid, gråhvid eller gullig, blød, slammet kalkbjergart (fig. 3), som næsten udelukkende består af kokkolitter, rester af µm­store plader fra marine alger (kokkosfærer). større dyrefossiler af snegle, mus­ linger, brachiopoder, søpindsvin og kiselsvampe træffes også. skrivekridtet er ofte gennemgravet af dyr, som har levet på havbunden eller inde i sedimentet (surlyk 1971, 2017; surlyk & håkansson 1999). skrivekridtet indeholder sort flint i mere eller mindre sammenhængende lag (fig. 3; madsen & stemmerik 2010). flintlagene viser stedvis, at skrivekridtet er af ­ lej ret på flade banker (surlyk et al. 2006) i et stort hav, som dækkede det meste af nordeuropa, og hvor der næsten ikke forekom klastisk sedimentation (fig. 4). dog træffes tynde, cyklisk aflejrede mergellag i dele af maastrichtien. skrivekridtaflejringerne fra sen kridt i østdanmark inddeles i to formationer: mandehoved formationen med to led (det meste af campanien) og møns klint formationen med fem led (sen campanien – ma ­ astrichtien, surlyk et al. 2013). inddelingen er baseret på under søgelser ved stevns klint (surlyk et al. 2006) og analyser fra en ca. 440 m dyb boring på stevns (stevns 1, fig. 5; surlyk et al. 2013). skrivekridtaflejringerne på møn er indpasset i denne litostratigrafi med møns klint formationen, som har fået navn efter klinten. typepro­ filet består af flere af de fem led. skrivekridtets alder er bestemt ved makro­ og mikrofossiler fra klinterne på høje møn og hvideklint. analyserne giver aldre fra nedre og øvre maastrichtien (surlyk 1984; surlyk & håkansson 1999; thomsen 1995; jelby et al. 2014). 50°n 60°n 0°ø 10°ø 20°ø 200 km møn n fig. 4 palæogeografisk kort over sen kridt i nordeuropa med fordeling af land (grønt) og hav (blåt). fra gravesen et al. (2017) efter damholt & surlyk (2012). fig. 4 palaeogeographic map of the late cretaceous sea in northern europe. green: land areas. blue: sea. from gravesen et al. (2017) after damholt & surlyk (2012). alder danien hvidskud led boesdal led flagbanke led højerup led sigerslev led rørdal led cerithium kalk led fiskeler ledn ø n ø ø v re k ri d t m aa st ic ht ie n m øn s kl in t fm m an de ho ve d fm c am pa ni en litostratigrafisk enhed fig. 5 stratigrafisk inddeling af øverste kridt og nederste danien i østdanmark (efter surlyk et al. 2006, 2013). fig. 5 stratigraphic division of the uppermost cretaceous and lowermost danian in eastern denmark (after surlyk et al. 2006, 2013). http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 7 of 53 www.geusbul let in.org glaciale og interglaciale aflejringer data og tolkning på prækvartæroverfladen ligger en op til 50–100 m tyk lagserie af glaciale sedimenter, der fortrinsvis er aflejret i den sidste del af weichsel­istiden (ca. 75 000 til 11 700 år før nu), men med få spor af ældre lag fra saale­istiden (390  000 til 130  000 år før nu) og eem­mellemistiden (130 000–115 000 år før nu). de glaciale sedimenter er på det meste af møn glacialtektonisk deformeret, foldet og forkastet, og ligeledes er store skiver af skrivekridt flyttet op fra den prækvartære overflade under møn eller fra havområdet omkring møn (f.eks. berthelsen 1979; pe dersen 2000). disse skrivekridtlag kan således betragtes som en del af de kvartære aflejringer. desuden indgår mindre slirer af skrivekridt i de klastiske aflejringer. under karteringen skelnes mellem følgende gla­ ciale jordarter: moræneler (ml), morænesand (ms) og morænegrus (mg). hertil kommer de glaciofluviale og glaciolacustrine jordarter smeltevandsgrus (dg), sand ler silt fi nt g ro ft g ru s st en bl ok ke ler og silt tørv og gytje sand sand, grus og sten moræneler diamikton flint glacitektonit skriveskridt lamineret strukturløst ribber klatrende ribber trug krydslejring planær krydslejring skarp laggrænse ler med sten gradvis laggrænse erosiv eller glaciotektonisk laggrænse holocæn eem saale kridt w ei ch se l holocæne a�ejringer litologi strukturer alder 10 m senglaciale ferskvandslag lolland till formation ny borre formation midtdanske till formation klintholm till formation kraneled formation ristinge klint till formation eem a�ejringer saale moræneler maastrichtien skrivekridt tidlig weichsel a�ejringer stubberup have formation kobbelgård formation litologi strukturer m el le m a fig. 6 a: skematisk strati gra fisk log af kvartære interglaciale, glaci ale og holocæne enheder repræsenteret på møn­kortbladet. b: beliggenheden af geografiske lokaliteter med reference til geologiske typelokaliteter. fig. 6 a: schematic stratigraphic log of the quaternary interglacial, glacial and holocene deposits occurring in the møn map sheet area. b: locations of geographic names with reference to geological type localities. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 8 of 53 www.geusbul let in.org smeltevandssand (ds), smeltevandssilt (di) og smelt­ evandsler (dl), og senglaciale jordarter: ferskvands­ grus (fg), ferskvandssand (fs), ferskvandsler (fl) samt skrivekridt (sk). en samlet oversigt over jordarter og symboler kan ses i pedersen et al. (2015), appendiks 2. den kvartærgeologiske stratigrafi er baseret på litte­ ratur, data og undersøgelser fra kystprofiler langs møns syd­, øst­ og nordkyster og fra få profiler i grusgrave samt på karteringen (fig. 6a). desuden udgør tvær­ profiler vist på det geologiske kort et væsentligt bidrag til forståelsen af den glaciale stratigrafi. en korrelation af formationerne med tværprofilerne er optegnet på fig. 25. ved udarbejdelsen af profilerne er anvendt boredata fra geus borearkiv og jupiter database (gravesen & fredericia 1984). et oversigtskort over vigtige geologiske lokaliteter på møn kan ses på fig. 6b. ældre kvartære aflejringer: saale, eem, tidlig weichsel og tidlig mellem weichsel (mis 6, 5e, 5a-5d, 4) spredte forekomster af aflejringer fra saale og eem, marine isotope stage (mis) 6 og yngre, er truffet i klin­ terne (fig. 6a). de indgår i de stærkt deformerede gla­ cialtektoniske enheder som fragmenter og flager. selv om de således ikke findes faststående, er det sand­ synligt, at de har været udbredt på dele af møn. saale (mis 6) lag af moræneler under eem­lagene ved kraneled, hellehavn nakke og stubberup have (de to sidstnævnte lokaliteter med hver to lag) betragtes som værende fra saale (fig. 8; hyde 1986). en del andre lokaliteter inde­ holder formodentlig også saale­moræneler, selv om den omfattende glacialtektoniske påvirkning kan gøre tolk­ ningen usikker (puggaard 1851; hintze 1937; hyde 1986; houmark­nielsen 1994). lagene af moræneler er min­ dre end 3 m tykke (fig. 7) og er antagelig fra sen saale (warthe). den glacialtektoniske påvirkning og deforma­ tion af både saale­ og eem­lagene skete i mellem eller sen weichsel (berthelsen 1973; aber 1979; pedersen & gravesen 2009; houmark­nielsen 2010). eem (mis 5e) danmark var i eem­mellemistiden omgivet af hav, og det antages, at eem­havet nåede ud over dele af møn (seidenkrantz et al. 2000), selv om en præcis udbred­ else er vanskelig at fastlægge på grund af de senere stubberup have hellehavns nakke liselund pomlerende taleren abildgaards fald dronningestolen gråryg siesø bjerg nælderende jættebrink kraneled klint kobbelgård klint tøvelde klint møn hjelm bugt stege bugt 12°30’ø hvid ek lin t stege hjelm nakke madses klint præstebjerg klintholm klint aborrebjerg ny borre 5 km 55°n nyord ulvshale kostervig ulvsund steg e n or b fig. 6 (fortsættes) a: skematisk strati gra­ fisk log af kvartære interglaciale, glaci ale og holocæne enheder repræsenteret på møn­kortbladet. b: beliggenheden af geografiske lokaliteter med reference til geologiske typelokaliteter. fig. 6 (continued) a: schematic stratigraphic log of the quaternary interglacial, glacial and holocene deposits occurring in the møn map sheet area. b: locations of geographic names with reference to geological type localities. weichsel moræneler, lolland till formation weichsel smeltevandssand weichsel moræneler, midtdanske till formation weichsel, stenfattigt ler eem marint ler saale smeltevandssand saale moræneler maastrichtien skrivekridt sv nø 5 m ? ? ? a fig. 7 eem­aflejringer ved hjelm nakke. a: feltskitse af saale­, eem­ og weichsel­lagene ved hjelm bugt fra berthelsen et al. (1977). b: smeltevandsler og ­sand (af saale­alder?) overlejret af marint eem­ler. foto: tove stockmarr (1996). fig.7. eemian deposits at hjelm nakke. a: field sketch of the saalian, eemian and weichselian deposits after berthelsen et al. (1977). b: clay and glaciofluvial sand (saalian?) overlain by marine eemian clay. photo: tove stockmarr (1996). http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 9 of 53 www.geusbul let in.org glaciotektoniske påvirkninger af aflejringerne. eem­ havet er påvist fra nordsøen gennem de danske far­ vande til langt ind i østersøen mod øst, med formodet passage i karelen i finland og videre til hvidehavet mod nordøst i tidlig eem (knudsen et al. 2011). selve den tidlige eem­transgression i østersøen ser ud til at være sket fra øst, med de ældste marine aflejringer i den fin­ ske bugt og yngre aflejringer i østdanmark, holland og tyskland (kristensen et al. 2000). den dominerende aflejring på møn er marint, grøn­ ligt til grønsort ler (cyprina ler), som skønsmæssigt er op til 2 m tykt (fig. 7), men de fleste steder kun omkring 0,5 m. de glacialtektoniske påvirkninger gør imidlertid denne vurdering usikker. marine fossiler er fundet i sand­ og lerlag på nord­ kysten ved øksnehøj/orebjerg, ved pomlerende og stub berup have og flere steder på møns klint, især i faldene, men også på ryggenes sider, f.eks. ved hun­ devangsfald, hvidskudsnakke, lille stejleberg, store stejlebjerg, grårygfald, fruerstuefald, nælderende, sandfaldet, skrædderrenden, maglevandsklinten/faldet, dronningestolen, vitmunds nakke, søndre hylledals­ klint, taleren og jydeleje fald (puggaard 1851; madsen et al. 1908; ødum 1933; hintze 1937; hansen & nielsen 1960; hyde 1986). desuden findes eem­aflejringer i kraneled klint øst for klintholm på møns sydkyst (hou­ mark­nielsen 1994) og på to lokaliteter i klinten mellem hjem nakke og hvideklint (fig. 7; berthelsen et al. 1977; stockmarr 1996). der er desuden under de marine lag i bl.a. store stejlebjerg fundet gytje og sand med fersk­ vandsmollusker og plantefossiler fra en mellemistid (n. hartz i hintze 1937). denne eem­lagfølge fra fersk­ vands­ til marine lag kan genfindes på andre lokaliteter i østersøen (kristensen et al. 2000; knudsen et al. 2011). makrofossiler er både fundet i fint sand og fedt ler. marine mollusker bærer ofte præg af påvirkning fra de glacialtektoniske episoder, men ligger øjensynligt in situ i leraflejringerne. antallet af arter (ødum 1933; hintze 1937) er beskedent i forhold til den omfattende liste over marine eem­fossiler i madsen et al. (1908). de fossilholdige lags litologi, fossiler og alder har været diskuteret, men undersøgelser af foraminiferer har demonstreret en marin eem­fauna ved hjem bugt, kraneled og pomlerende (v. madsen i hintze 1937; ber thelsen et al. 1977; j. frederiksen i hyde 1986; p. kristensen i houmark­nielsen 1994), hvilket således underbygger tilstedeværelsen af makrofossiler fra et eem­hav. stedvis optræder også ler og sand mellem to lag af moræneler, hvor alle lag indeholder foraminiferer fra eem, hvilket viser, at disse lag er omlejret (konradi 1973). foraminifererne i eem­aflejringerne har vist sig at til­ høre en boreo­lusitansk, kystnær fladvandsfauna (ber­ thelsen et al. 1977), som kan sammenlignes med faunaer fra andre lokaliteter i den vestlige del af østersøen og i nordsøen (konradi 1976; knudsen 1991), som har levet i havvand med lidt højere temperaturer og saltholdighed end i dag (houmark­nielsen 1994). dateringer af eem­aflejringer fra klintholm–kraneled ved hjælp af optisk stimuleret luminiscens (osl) og termo­ luminiscens (tl) har givet aldre, der er væsentligt yngre end forventet for eem (yngre end 115 000 år før nu). en nærmere diskussion af eventuelle årsager kan ses i hou­ mark­nielsen (1994, 2008, 2010). aminosyredateringer på udvalgte mollusker (arctica islandica, turritella communis og nassa reticulata fra klintholm, arctica islandica og turritella communis fra stubberup have og turritella communis fra hjelm nakke (petersen 1984; miller & mangerud 1985; houmark­nielsen 1994, 2008)) understøtter, at aflej ringerne fra disse klintprofiler er af eem­alder. fig. 7 (fortsættes) eem­aflejringer ved hjelm nakke. a: feltskitse af saale­, eem­ og weichsel­lagene ved hjelm bugt fra berth­ elsen et al. (1977). b: smeltevandsler og ­sand (af saale­alder?) overlejret af marint eem­ler. foto: tove stockmarr (1996). fig.7. (continued) eemian deposits at hjelm nakke. a: field sketch of the saalian, eemian and weichselian deposits after berthelsen et al. (1977). b: clay and glaciofluvial sand (saalian?) overlain by marine eemian clay. photo: tove stockmarr (1996). b http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 10 of 53 www.geusbul let in.org tidlig weichsel og tidlig mellem weichsel (mis 5a-5d, mis 4) i tidlig og tidlig mellem weichsel var møn formodentlig isfrit og forholdene domineret af jordflydning og trans­ port og aflejring i floder, dog med begrænset udbred­ else. omlejrede eem­makrofossiler findes i disse lag (houmark­nielsen 1994), der vurderes til at være mel­ lem 2 og 20 m tykke. mellem og sen weichsel (mis 3 og 2) de glaciale enheder, der har den største udbredelse inden for kortbladet, er aflejret i mellem weichsel (for 74 000– 29 000 år siden) i den sidste del af den seneste istid, og i sen weichsel (for 29 000 til ca. 11 700 år siden, fig. 6a). enhed­ erne repræsenterer kun dele af de nævnte tidsafsnit. på trods af en god blotningsgrad er en opstilling og beskrivelse af nye formelle litostratigrafiske enheder og beskrivelse af de eksisterende samt påvisning af deres udbredelse en vanskelig opgave på grund af de omfattende glacialtektoniske forstyrrelser af aflejring­ erne. imidlertid er de geologiske profiler med litologi, dannelse og alder i en del af klinterne velbeskrevet i videnskabelige publikationer fra de senere år, f.eks. houmark­nielsen (1994); gravesen et al. (2017), og i rap­ porter og upublicerede specialeafhandlinger. desuden har hydrogeologisk kortlægning (rambøll 2006, 2007) bidraget til forståelsen af den geologiske opbygning og lagenes rumlige udbredelse i det indre af møn, hvor der kun er få daglokaliteter. disse informationer i kombi­ nation med kartering og undersøgelse af kystprofiler, profiler i råstofgrave, boredata og geofysik har bidraget til at underbygge opstillingen eller genbeskrivelsen af formationer fra mellem og sen weichsel. alle alders­ angivelser er baseret på houmark­nielsen et al. (2017). ud fra udvalgte logs fra kystklinterne er der udarbej­ det et længdeprofil, som kan følges fra madses klint langs møns sydøst­ og sydkyst til østkysten langs møns klint og videre mod nordøst til hellehavn nakke for at lolland till fm kobbelgård fm klintholm till fm kraneled fm ristinge klint till fm tidlig weichsel saale moræneler senglaciale ferskvandslag ny borre fm midtdanske till f m stubberup have fm 0 60 km madse s k lin t hjelm n akke tøvelde klin t nord lig e m øns k lin t helle hav ns n akke stu bberu p h ave bru nsh av pomlerende klin th olm lise lund hvide klin t 10 20 28 32 33 36 39 5 10 15 20 25 30 36 vest øst vest stubberup have fm midtdanske till fm ny borre fm lolland till fm m nordøst v ø v nø fig. 8 sammensat længdeprofil med tykkelser og horisontal udbredelse af de litostratigrafiske enheder på møns syd­, øst­ og nord­ kystklinter baseret på udvalgte vertikale logs. indekskortet viser, hvor profilet ændrer retning. bemærk forskellen på vertikal og horisontal skala. baseret på houmark­nielsen (2010) og gravesen et al. (2017). fig.8. composite section showing the thicknesses and horizontal distribution of the lithostratigraphical units along the south, east and north coast cliffs of møn based on selected vertical logs. the index map shows where the section changes orientation. note different horizontal and vertical scales. based on houmark-nielsen (2010) and gravesen et al. (2017). http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 11 of 53 www.geusbul let in.org slutte på nordkysten ved brunshoved (fig. 8). profilet viser den samlede litostratigrafiske opbygning med de forskellige enheders vertikale og horisontale udbre­ delse. de litologiske enheder og deres opdeling i forma­ tioner er beskrevet nedenfor. ristinge klint till formation (mis 3) navn og historie. formationen har sit navn efter ristinge klint på langeland. det er en klassisk lokalitet, der har bidraget til forståelsen af danmarks geologiske opbyg­ ning (madsen et al. 1908; madsen 1916). andersen (1945) og ber thelsen (1973) kaldte morænelerenheden for den gammel baltiske till. houmark­nielsen (1987) opstillede formelt formationen og gav en detaljeret beskrivelse af dens dannelseshistorie. han påviste blandt andet, at moræneleret, der indeholder fingrus med en baltisk proveniens, blev aflejret i forbindelse med et isfremstød fra sydøst (sjørring et al. 1982). på møn er formationen ved klintholm blevet betegnet som ”grey­red till, unit 4” af houmark­nielsen (1994). typelokalitet. typelokaliteten er ristinge klint på lange­ land (hou mark­nielsen 1987). litologi. formationen består af moræneler med en grå eller rødgrå farve, og den rødlige farve skyldes et højt indhold af baltiske sandsten og kalksten fra nedre palæozoikum. dens nedre del består ofte af båndet, gråt og hvidligt, glacialtektonisk udvalset moræneler med et højt indhold af opbrudt, shearet og foldet skrivekridt (fig. 9; banham 1977; pedersen 1988). den øvre del af moræneleret er rødgråt med rødligt sand og sandsten, som ofte kan findes på stranden foran klinterne. et indhold af fingrus og klaster viser en baltisk dominans (berthelsen et al. 1977; hyde 1986; houmark­nielsen 1994; stockmarr 1996; smed 2010). i formationen findes undertiden omlejret eem­materiale. på det nordøst­ lige møn er der i klintprofilerne målt en till fabric, som viser en bred spredning omkring en østlig retning med et svagt til moderat dyk mod øst (50 målinger i hyde 1986). bortset fra, at moræneleret også her er påvirket af senere glacialtektonisk dislokation, viser målingerne en veldefineret isstrømsretning fra øst mod vest. denne proveniens understøttes af fingrustællingen med 25 % kalksten, hvoraf 2/3 er palæozoiske, et lavt flintindhold (<5 %) og op til 15 % sedimentære gruskorn, overvejende af sandsten (hyde 1986). jordartssymboler: ml, ms. grænser og tykkelse. ved klintholm er den nedre grænse en glacialtektonisk inkonformitet på ældre, formodent­ lig tidlig eller tidlig mellem weichsel­aflejringer. ved hvideklint og hjelm nakke hviler formationen ofte på skrivekridt. den øvre grænse mod kraneled forma­ tionen er sedimentær og erosiv. formationens tykkelse er antagelig få meter ved hvideklint – hjelm nakke, 4 m ved klinthavn, 1 m ved høje møn, 3 m ved hellehavn nakke og 2 m ved pomle rende (hyde 1986; houmark­nielsen 1994; stockmarr 1996). udbredelse og korrelation. formationen er truffet på en række lokaliteter som ved hjelm nakke – hvide­ klint (antagelig), høje møn, klintholm, hellehavn nakke og pomlerende, og dette antyder, at formationen har en større udbredelse på det sydlige og østlige møn. midtdanske till fm kobbelgård fm risti nge klin t t ill f m sk riv ek rid t fig. 9 rødbrunt moræneler tilhørende ristinge klint till forma­ tionen ved store stejlebjerg. formationen hviler på en inkon­ formitet på toppen af skrivekridt fra møns klint formationen. her i den sydlige del af møns klint overlejres ristinge klint till formationen af ler og fint sand tilhørende kobbelgård forma­ tionen. på toppen af lagserien findes moræneler fra den midt­ danske till formation. landmålerstokkens enheder er 20 cm brede. fig. 9. red-brown clayey till of the ristinge klint till formation at store stejlebjerg, møns klint. the formation rests unconformably on top of the cretaceous chalk of the møns klint formation. finegrained lithologies of the kobbelgård formation overlie the ris tinge klint till formation. at the top, clayey till of the mid danish till formation is present. for scale, the units on the ranging pole are 20 cm wide. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 12 of 53 www.geusbul let in.org for mationen har antagelig udbredelse i store dele af danmark (houmark­nielsen 1987, 2007). alder. formationen er aflejret i forbindelse med ris­ tinge isfremstødet, som foregik for mellem ca. 55  000 og 50 000 år siden (houmark­nielsen 2010; houmark­ nielsen et al. 2017). aflejringsmiljø. moræneleret er dannet som en bundmoræneaflej ring. indholdet af baltisk fingrus og klaster samt de glacialtektoniske strukturer viser, at bevægelsen af glet sjeren skete fra sydøst mod nordvest (houmark­nielsen 1999; smed 2010). denne isstrøm kaldes for ristinge isfremstødet (det gammelbaltiske isfremstød, hou mark­nielsen 1987; houmark­nielsen et al. 2017). kraneled formation (mis 3) ny formation navn og historie. kraneled formationen har navn efter lokaliteten kraneled øst for klintholm på sydlige møn (fig. 6b). houmark­nielsen (1994) betegnede den som ”glaciofluvial and lacustrine deposits, unit 5”, mens den i hou mark­nielsen & kjær (2003) blev angivet som ”klintholm beds” og i houmark­nielsen et al. (2017) som ”klintholm”. typelokalitet – og profil. typelokaliteten er ved kraneled klint øst for klintholm (fig. 10; houmark­nielsen 1988, 1994). typeprofilet ses i houmark­nielsen (1994, fig. 5, unit 5). litologi. den nedre del består af krydslejret og imbrikeret sand og grus. den øvre del er opbygget af kryds lejret sand, som fortsætter i finkornet sand og horisontalt lamineret ler (fig. 10). der kan være få planterester. jordartssymboler: dl, ds, dg. kraneled fm ristinge klint till fm kalk glacitektonit skrivekridt fig. 10 finkornede smeltevandsaflejringer tilhørende kraneled formationen, hvilende på ristinge klint till formationen. en skrivekridt­glacitektonit ses ved formationens basis langs toppen af skrivekridtet. klinten ved kraneled, klintholm. fig. 10. fine-grained meltwater deposits of the kraneled formation resting on the ristinge klint till formation. between the till and the underlying chalk, a chalk-glacitectonite occurs. kraneled cliff, klintholm. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 13 of 53 www.geusbul let in.org grænser og tykkelse. den nedre grænse til ristinge klint till formationen er en skarp, erosiv grænse. den øvre grænse er en gradvis sedimentær overgang til klintholm till formationen. formationens tykkelse er omkring 4 m ved klintholm og 2 m ved tøvelde klint, men større på det vestlige møn (omkring 7 m i boringer). udbredelse og korrelation. blotningsgraden er ikke god, så det er vanskeligt at fastslå formationens udbredelse. den træffes i hjelm bugt og på vestmøn. undersøgelser ved boringer på kriegers flak i østersøen øst for møn viser en interstadial søaflejring fra et tidsrum (for 40 000–35 000 år siden) antagelig sva­ rende til kraneled formationen (anjar et al. 2010, 2012). samtidige aflejringer findes også inden for det store issøområde ved brorfelde, hvor der i tykke sandlag er fundet stødtænder af mammut, som er dateret til ca. 44 000 år før nu (selsing 1982). interstadiale aflejringer på sejerø er dateret til omkring 36 000 år før nu, hvilket placerer dem før klintholm till formationen og antagelig samtidig med lagene fra kraneled formationen. planteindholdet i de lakustrine aflejringer på sejerø peger på et åbent, træløst land med tundravegetation (kolstrup & hou mark­nielsen 1991; bennike et al. 2007). den finkornede del af kraneled for­ mationen kan repræsentere det samme aflejringsmiljø som på møn under sandnæs mildningen. alder. formationen har en alder på ca. 45  000–35  000 år før nu (houmark­nielsen et al. 2017) og er således formodentlig dannet under den første fase af sandnæs mildningen. aflejringsmiljø. de nederste, grovkornede lag er glacio­ fluviale aflejringer, som antageligt er afsat i en smelte­ vandsflod i et periglacialt område. fortsættelsen op ad peger på dannelse i et lakustrint aflejringsmiljø. landskabet må forventes at have været åbent med beskeden tundravegetation under arktiske forhold. klintholm till formation (mis 3) justeret formation navn og historie. formationen har navn efter den lille by klintholm på sydkysten af møn. den er beskrevet af houmark­nielsen (1994) som ”thick grey diamic­ ton” og af krüger og kjær (1999) som ”unit 6”. siden­ hen er till­ aflejringen blevet betegnet som klintholm till (hou mark­nielsen & kjær 2003; houmark­nielsen 2008, 2010; houmark­nielsen et al. 2016). typelokalitet og -profil. typelokaliteten og typeprofilet er kraneled øst for klintholm, mens et referenceprofil ses ved kobbelgård vest for klintholm (fig. 11; hou mark­ nielsen 1988, 1994, fig. 5, unit 6). litologi. klintholm till formationen består af gråt, mas­ sivt, siltet og sandet moræneler med et moderat indhold af spredte klaster (fig. 6, 11). øverst kan der optræde mørkebrunt, massivt moræneler, der stedvis er lagdelt. der træffes stedvis omlejrede skaller eller skalfrag­ menter af marine mollusker fra eem. jordartssymboler: ml, ms. grænser og tykkelse. formationens tykkelse er ca. 2,6 m ved madses klint, 3,6 m ved hjelm nakke – hvideklint, 4 m ved tøvelde klint, 3–5 m ved klintholm, 3 m ved møns klint, ca. 1,5 m ved hellehavn nakke og 1,8 m ved stubberup have (hyde 1986; houmark­nielsen 1994; stockmarr 1996; gravesen et al. 2017). klintholm till formationens nedre grænse er en grad­ vis overgang fra leret moræneler til sand og mudder i kraneled formationen, men denne grænse er under­ tiden også en glacialtektonisk inkonformitet. den øvre grænse er gradvis fra moræneler til mudder i kobbel­ gård formationen. udbredelse og korrelation. klintholm till formationen har stor udbredelse fra møn til falster (korselitse) og videre over sydsjælland (houmark­nielsen 2010). alder. klintholm till formationen er aflejret under klintholm isfremstødet, som har en alder på ca. 35 000– 32  000 år før nu (houmark­nielsen 2010; houmark­ nielsen et al. 2017). aflejringsmiljø. formationens morænelersaflejringer er afsat af en gletsjer, som har bevæget sig fra sydøst mod nordvest. dette kan demonstreres ved fingrustællinger og ’klast fabric’ orienteringsmålinger samt glacial­ tektoniske strukturer. smed (2010) har analyseret klastindholdet ved stentælling, og dette viser, at en del af gletsjerbevægelsen har været omtrent fra nord­ øst mod sydvest gennem sverige, hvorefter retningen ændredes til vestlig mod danmark (se også ukkonen et al. 2007). aflejringen anses for at være for at være sket langs bunden af gletsjeren (bundmoræneaflejring). isfremstødet betegnes af houmark­nielsen et al. (2017) som klintholm isstrømmen. kobbelgård formation (mis 3) ny formation navn og historie. kobbelgård formationen er navngivet efter kobbelgård klint, som ligger ved hjelm bugt på det sydlige møn, hvor der er truffet interstadiale lakustrine aflejringer samt glaciolakustrine og ­fluviale aflejringer. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 14 of 53 www.geusbul let in.org aflejringerne er især godt undersøgt omkring klintholm havn i de to klinter ved kobbelgård og kraneled (hou­ mark­nielsen 1988, 1994, 2003, 2010; krüger & kjær 1999; hou mark­nielsen et al. 2016), hvor en lagfølge, som også omfatter klintholm till og kraneled formationerne (fig. 8), er detaljeret beskrevet og tolket ud fra alders­ bestemmelser, sedi mentologi og indholdet af plante­ og dyrefossiler. den nedre del af leraflejringerne er kendt fra en række tidlige undersøgelser på møn (f.eks. puggaard 1851; hintze 1937; hansen & nielsen 1960; ber thelsen et al. 1977) som en leraflejring med et beskedent klastindhold (”stenfattigt ler”, hyde 1986). et karakteri­ stisk sprækkesystem (”søjleforkløftning”, berthelsen et al. 1977) har givet ophav til navnet ’elefantfodleret’ på grund af, at uderoderede og ofte foldede partier af leret har lighed med en elefants fødder. ved klintholm er lagene beskrevet som ”lower lake sediments, unit 7” af houmark­nielsen (1994) og som en del af ”upper kob­ belgård beds” i houmark­nielsen et al. (2016). den øvre del af formationen omfatter overvejende den fossilrige del af søaflejringerne ved klintholm, som består af ler, silt og diamikte aflejringer. de er beskrevet af houmark­nielsen (1994) som ”upper lake sediments, unit 8”, ”upper kobbelgård beds ” og ”kobbelgård beds” (houmark­nielsen & kjær 2003; houmark­nielsen 2010). aflejringerne er også beskrevet af krüger & kjær (1999). typelokalitet og profil. typelokaliteten for kobbelgård formationen er ved kobbelgård klint vest for klintholm havn og kraneled klint øst for, hvor der er opmålt flere delprofiler (se hou mark­nielsen 1994 fig. 5, unit 7, unit 8; houmark­nielsen 1988, 1994; krüger & kjær 1999). litologi. formationen består nederst af ler/mudder med tynde sandlag og mange spredte klaster (fig. 12), og ler­ aflejringen er gennemsat af talrige sprækker. ved basis findes stedvis krydslejret og horisontalt lagdelt sand, undertiden med omlejrede skaller fra eem­aflejringer. leret er fedt og siltet, kalkholdigt og mørkegråt, med tynde sandlag og ofte spredte klaster. indholdet af silt og ler er 90–94 %, resten er sand og fint grus (hyde 1986). fingrusindholdet er helt domineret af palæo­ zoiske kalksten (hyde 1986). foruden sprækkesystemer er der kun fundet strukturer i form af en vanskeligt gen­ kendelig og utydelig lagdeling i leraflejringerne. der er spredte planteaflejringer i form af blade og plantedetri­ tus. desuden er der fundet pollen fra en arktisk fauna (kolstrup & houmark­nielsen 1991). formationens øvre del består af en række vekslende, massive ler­ og siltlag samt diskontinuerlige lag af finkor­ net sand med indslag af lerede diamikte lag mod top­ pen (fig. 13). lagene indeholder fragmenter af planter inkl. blade og finkornet detritus, og der er beskrevet en omfattende terrestrisk mikro­ og makroflora (kolstrup & houmark­nielsen 1991; bennike et al. 1994, 2007). klaster og fingrus i de diamikte aflejringer er domineret af palæozoiske kalksten og skifre. de øverste lag inde­ holder imidlertid en blanding af lokale og baltiske bjerg­ artsfragmenter. jordartssymboler: dl, di, ds, ml. grænser og tykkelse. formationens nedre grænse til klitholm till formationen er oftest gradvis. den øvre grænse til den midtdanske till formation er en glacial­ tektonisk inkonformitet. kobbelgård formationens tykkelse kan nå helt op på 15 m ved klintholm, men varierer ved andre klinter med følgende ca.­værdier: madses klint: 7,5 m, hjelm nakke – hvideklint: 12,5 m, tøvelde klint: 7,5 m, møns klint: 12 m, liselund: 5 m, hellehavn nakke: 11 m, og stubberup have: 7 m (ødum 1933; hintze 1937; hyde 1986; houmark­nielsen 1988, 1994; stockmarr 1996; hou mark­nielsen et al. 2016; gravesen et al. 2017). lolland till fm kobbelgård fm klintholm till fm ristinge klint till fm skrivekridt fig. 11 gråt moræneler tilhørende klintholm till formationen i klinten ved kraneled, klintholm. i profilet ses desuden aflej­ ringer af skrivekridt og ristinge klint till, kobbelgård og lolland till formationerne. fig. 11. grey clayey till of the klintholm till formation at kraneled cliff, klintholm. the section also contains deposits of chalk as well as deposits of the ristinge klint till, kobbelgård and lolland till formations. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 15 of 53 www.geusbul let in.org udbredelse og korrelation. formationen har stor udbred­ else på det sydlige, østlige og nordøstlige møn. den øvre fossilrige del træffes især på sydkysten. dele af forma­ tionen er sandsynligvis samtidige med aflejringer på kriegers flak (anjar et al. 2010, 2012). desuden findes formationen langs falsters østkyst, bl.a. ved korselitse (houmark­nielsen 2010; gravesen et al. 2017). alder. aflejringerne er fra mellem weichsel, mis 3. for­ mationen hviler på klintholm till formationen, som har en alder på 35 000–32 000 år før nu (houmark­nielsen et al. 2017). der er foretaget en række 14c­, osl­ og tl­aldersbestem­ melser på plantemateriale og sedimenter fra kobbelgård formationen, som giver en spredning over den sidste del af mellem weichsel. formationen er således dateret til 35 000–20 000 år ved osl­metoden og mellem 37 000– 25 000 år ved 14c­metoden (houmark­nielsen 2010). dateringerne fra klintholm­området antyder tilsam­ men, at det sandsynlige aldersinterval ligger mellem 32  000 og 28  000 år før nu (houmark­nielsen & kjær 2003; houmark­nielsen 2008, 2010; houmark­nielsen et al. 2016; houmark­nielsen et al. 2017). formationen har en alder på ca. 32 000–28 000 år og repræsenterer en lakustrin aflejring fra den sidste del af sandnæs mildningen (houmark­nielsen et al. 2017). aflejringsmiljø. ved slutningen af den kolde periode for a b fig. 12 lakustrine aflejringer af gråt, siltholdigt ler med få gruskorn og sten tilhørende kobbelgård formationen. a: typisk glaciolakustrint ler med få spredte dropsten. klinten ved liselund. b: ler ved foden af jordskredskomplek­ set neden for liselund. fig. 12. lacustrine deposits of grey, silty clay with scattered pebbles of the kobbelgård formation. a: typical glaciolacustrine clay with widely spaced dropstones. b: clay at liselund in the dislocation horizon at the toe of a landslide. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 16 of 53 www.geusbul let in.org ca. 32  000 år siden, da klintholm isfremstødet, som aflejrede klitholm till formationen, smeltede væk, blev der dannet en stor issø, der var udbredt over hele den sydlige, østlige og nordlige del af møn og måske endog over et større område omkring østersøen, hvilket også er blevet foreslået af hansen (1965) og houmark­niel­ sen (1994). i starten blev der aflejret sand og grus, hvor smeltevand strømmede ind i søen. derefter var der udstrakt aflejring af ler og silt, og isflager bragte klast­ materiale ud i søen, der blev aflejret som dropsten. i det vegetationsløse landskab blev finkornet sand blæst ud i søen og aflejret i tynde lag. i takt med, at klimaet blev varmere, og tilførsel af smeltevand fra den baltiske is blev formindsket, blev søen gradvis mindre. efter hånden bredte en arktisk vegetation sig omkring søen med græs og urter samt enkelte små træer som fyr og birk. i den sidste del af perioden for ca. 30 000–28 000 år siden var søen sandsynligvis indskrænket til at dække et område omkring klintholm. mudder (ler, silt og fint sand) blev aflejret fra suspension, mens de diamikte lag er dannet ved mudderstrømme, der er gledet ud i søen, eller som stammer fra isflager. også vindblæst fint sand og silt er påvist (houmark­nielsen 1994; krüger & kjær1999). vegetationen omkring søen var sub arktisk uden træer, men med urter og mange buske. det stadigt varmere subarktiske klima medvirkede til, at plante vækst bredte sig i og omkring søen. flodsyste­ mer omkring søen aflejrede lateralt herfor krydslejrede, finkornede banker, og floder førte sandet og gruset materiale ud i søen (stubberup have formationen). fund af en mammuttand ved slotshøj, stege nor (aaris­sørensen et al. 1990) kan indikere tilstedeværelsen af en større græsslette omkring søen. på dette tidspunkt har større græs­ og urtestepper med mammutter og andre større dyr eksisteret i danmark (aaris­sørensen et al. 1990) og det sydlige sverige (berglund et al. 1976; ukkonen et al. 2007). fig. 13 gråt, siltet og fedt, lagdelt ler tilhørende kobbelgård formationen. klinten ved kobbelgård, klintholm. landmålerstokkens enheder er 20 cm brede. fig. 13. grey, silty and fine-grained, horizontally laminated clay of the kobbelgård formation. kobbelgård cliff, klintholm. for scale, the units on the ranging pole are 20 cm wide. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 17 of 53 www.geusbul let in.org stubberup have formation (mis 3) ny formation navn og historie. formationen har navn efter den nord­ lige del af møns klint fra området ved stubberup have. typelokalitet og -profil. typeprofil og ­lokalitet er stubbe­ rup have (fig. 14; hyde 1986). litologi. formationen består af sand og grus samt under­ tiden sten (fig. 8, 14). sedimentære strukturer forekom­ mer som ribber og storskala krydslejring, horisontal lagdeling og grovkornede kanaler med grus og sten, som skærer ned i de underliggende sandlag. finkornet, vel­ sorteret sand og silt med klatrende ribber træffes også. grænser og tykkelse. tykkelsen af stubberup have for­ mationen varierer en hel del (fig. 8). følgende er regi­ streret i klinterne: madses klint: 1,3 m, hjelm nakke – hvideklint: 4,5 m, tøvelde klint: 5 m, møns klint: 1,4 m, liselund: 10 m, hellehavn nakke: 11 m og pomle rende: a ler og silt sand sand sand, grus og sten litologi lamineret krydslejring strukturer skarp laggrænse ? moræneler ler silt �n m el le m gr ov g ru s og st en 2 m 1 2 3 0 4 5 8 7 6 11 10 9 12 midtdanske till formation stubberup have formation kobbelgård formation b fig. 14 a: mellem­ og grovkornet, krydslejret sand tilhørende stubberup have formationen. klinten ved stubberup have, type lokalitet for formationen. landmålerstokkens enheder er 20 cm brede. b: sedimentologisk log for stubberup have formationen. fig. 14. a: mediumand coarse-grained, cross-bedded sand of the stubberup have formation at the cliff of stubberup have, the type locality for the formation. the units on the ranging pole are 20 cm wide. b: sedimentological log of the stubberup have formation. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 18 of 53 www.geusbul let in.org mindst 3 m (hintze 1937; hyde 1986; stockmarr 1996; gravesen et al. 2017). enkelte boringer peger på en tyk­ kelse på op til 20 m, hvor lagene dog kan være glacial­ tektonisk forstyrrede. formationen har en skarp, nedre sedimentær grænse til kobbelgård formationen (fig. 8). mod vest fra tøvelde og klintholm klinterne til madses klint er grænsen til kobbelgård formationen formodentlig sammenflettet. den øvre grænse til den midtdanske till formation er er en skarp glacialtektonisk inkonformitet. udbredelse og korrelation. stubberup have formationen er udbredt på store dele af møn. alder. stubberup have formationens alder er forment­ lig delvis den samme som kobbelgård formationens (32 000–28 000 år; houmark­nielsen et al. 2017). aflejringsmiljø. aflejringsmiljøet var en slette med flettede floder, som har dannet små ribber og større banker. der har også stedvis været lakustrine forhold med aflejring af fint sand og silt under stor materialetilførsel. fig. 15 gråt moræneler tilhørende den midtdanske till formation. madses klint. fig. 15. grey clayey till of the mid danish till formation, madses klint. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 19 of 53 www.geusbul let in.org midtdanske till formation (mis 2) navn og historie. den midtdanske till formation (mid danish till formation) blev opstillet af houmark­niel­ sen (1987) for den litologiske enhed af moræneler, der dækker danmark fra landets østlige grænse mod vest til hovedopholdslinjen i den centrale del af jylland. for­ mationen har typelokalitet i ristinge klint på langeland, hvor den danner en ca. 15 m tyk, homogen litologisk enhed i de flager, som indgår i det glacialtektoniske kompleks i ristinge klint. baggrunden for opstillingen af formationen findes i houmark­nielsen (1987). typelokalitet og referenceprofil. typelokalitet og ­profil er ristinge klint på langeland (houmark­nielsen 1987). referenceprofilet på møn er ved madses klint (fig. 15). litologi. formationen består af massivt, kalkholdigt moræneler. der er ofte et højt indhold af kalkklaster. kalkindholdet i de glacialtektonisk påvirkede dele af moræneleret klassificerer disse dele som glacitektoniter (pedersen et al. 2018). jordartssymbol: ml. grænser og tykkelse. midtdanske till formation har på møn en tykkelse på op til 10 m, men tykkelsen varierer i øvrigt meget. i klinterne træffes følgende tykkelser: mad­ ses klint: 3,5 m, hjelm nakke – hvideklint: 6 m, tøvelde: 4 m, møns klint: 9 m, liselund: 3 m, pomlerende: varie­ rende tykkelse, stubberup have: 5,5 m og brunshoved: 8 m. (ødum 1933; hansen & nielsen 1960; berthelsen et al. 1977; hyde 1986; stockmarr 1996; gravesen et al. 2017). formationen overlejrer stubberup have forma­ tionen med en skarp grænse, almindeligvis med en gla­ cialtektonisk inkonformitet (fig. 6), og overlejres af ny borre formationen med en skarp sedimentær grænse. udbredelse og korrelation. formationen findes på hele møn undtagen i klint holm­området. alder. alderen er ca. 23 000–20 000 år før nu (hou mark­ nielsen et al. 2017). formationen er aflejret under det sidste glaciale maksimum i weichsel (’late glacial maxi­ mum’, lgm). aflejringsmiljø. den midtdanske till formation er en bundmoræne (lodgement till) aflejret af en gletsjer. glacialtektoniske retningsindikatorer, klastindhold og fingrustællinger viser, at gletsjeren har bevæget sig fra nordøst mod sydvest (berthelsen et al. 1977; aber 1979; hyde 1986; stockmarr 1996; smed 2010). formationen er aflejret under den maksimale udbred else af den skandinaviske iskappe i sidste del af weichsel­istiden. hovedopholdslinjen for dette isfrem­ stød var beliggende 100 km syd og vest for møn i et strøg fra flensborg og nord om hamburg til branden­ burg, hvorfra den drejede syd om berlin og ind i det nordlige polen, den såkaldte brandenburg israndslinje (smed & ehlers 2002; ehlers et al. 2004). ny borre formation (mis 2) ny formation navn og historie. formationen er opkaldt efter landsbyen ny borre syd for borre på østmøn. formationen inklu­ derer de overfladenære glaciofluviale og ­lakustrine sand­ og grusaflejringer, som kan træffes over store dele af møn. disse lag har ikke tidligere været navngivet. der indvindes sand og grus fra formationen i bl.a. ny borre grusgrav. typelokalitet og -profil. typelokaliteten findes i den aktive grusgrav i ny borre. typeprofilet kan ses på fig. 16, der viser en lagfølge af sand med indhold af både grov­ og finkornede aflejringer. litologi. formationen består af lagdelte silt­, sand­ og gruslag med stor udbredelse på møn. kornstørrelsen i formationen kan variere en hel del alt efter den sam­ menhæng, som lagene findes i. den nederste del består af finkornet og mellemkornet sand samt silt i vekslende lag med linser af grovere sand. undertiden indgår lerlag. lagene hælder stedvis og kan være dele af kanaler eller større, hældende forsæt, som viser en strømretning ler og silt sand sand litologi lamineret strukturer sand, grus og sten ribber trug krydslejring skarp laggrænse strukturløst moræneler ler silt �n m el le m gr ov g ru s og st en 2 m fig. 16 sedimentologisk log af ny borre formationens fluviale– lakustrine sekvens på typelokaliteten i ny borre grusgrav. fig. 16. sedimentological log of the fluvial–lacustrine sequence of the ny borre formation at the type locality in the ny borre gravel pit. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 20 of 53 www.geusbul let in.org fra syd og sydvest (fig. 17). den øvre del, som har størst udbredelse, består af krydslejret, grovere sand med gruslag, som forefindes i større kanalstrukturer. jordartssymboler: ds, dg, di, fs, fg. grænser og tykkelse. formationens nedre grænse mod den midtdanske till formation er skarp og overvejende sedimentær. den kan ses ved de samme lokaliteter som den øvre grænse. formationens øvre grænse til lolland till formationen er ligeledes skarp og ofte en glacialtektonisk inkonformi­ tet. grænsen kan træffes i en række kystklinter, bl.a. i madses klint, hjelm nakke, hvideklint, tøvelde, møns klint og i klinterne nord for høje møn (liselund, stubbe­ rup have og pomlerende). formationens tykkelse varie­ rer 5–10 m på store dele af øen, men kan være tykkere i de store kanaler. på typelokaliteten i ny borre grusgrav er tykkelsen ca. 6 m (storstrøms amt 2004; region sjæl­ land 2012). variationen af formationens tykkelse i klin­ terne, hvor den er overlejret af lolland till formationen, er: madses klint 12 m, hvide klint: 6 m, tøvelde: 1,5 m, møns klint: 5,5 m, liselund: mindst 4,5 m, pomlerende: op til 3 m og stubberup have: 5,5 m (hintze 1937; hyde 1986; stockmarr 1996). fig. 17 hældende, vekslende silt­, sand­ og gruslag i en lakustrin lagfølge fra ny borre formationen. ny borre grusgrav. fig. 17. inclined, alternating silt, sand and gravel layers in a lacustrine sequence of the ny borre formation. ny borre gravel pit. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 21 of 53 www.geusbul let in.org udbredelse og korrelation. formationen er udbredt på store dele af møn og ses på det geologiske kort i over­ fladen eller lige under lolland till formationen. dette kan også ses i rambøll (2006). formationen følger her en række dalstrukturer fra nord til syd omfattende borre sømose, maglemose, klostervig lavningen omkring stege nor, røddinge lavningen og fanefjord, som alle er tunneldale eller afvandingskanaler fra smeltningen af den ungbaltiske is. i en formodet tunneldal fra kob­ belgård til stege nor træffes råbylille ås, og i fanefjord tunneldalen findes fanefjord ås (smed 2014). i begge tunneldale træffes en del af formationens sedimenter. i hatformede bakker på det sydvestlige møn træffes for­ mationen også (berthelsen 1980). alder. ny borre formationen anses for at være aflej­ ret under fremrykningen af den ungbaltiske is. disse begivenheder antages at have foregået for 19  000– 18 000 år siden (houmark­nielsen et al. 2017). aflejringsmiljø. formationen repræsenterer flere forskel­ lige successive aflejringsmiljøer dannet af smeltevand fra den ungbaltiske is, nemlig et glaciofluvialt aflejrings­ miljø, hvor sand og grus aflejredes på dal­sandurer, der udvikledes i en række store afvandingskanaler eroderet ned i den midtdanske till formation, og et glaciolacus­ trint miljø, hvor smeltevandsaflejringerne blev aflejret i mere eller mindre isdæmmede langsøer vinkelret på isranden. sand­ og gruslag er også aflejret i overflade­ nære åse og hatformede bakker. lolland till formation (mis 2) navn og historie. lolland till formation er opkaldt efter lolland i det sydøstlige danmark og blev opstillet i forbindelse med beskrivelsen af kortbladet sakskøbing (pedersen et al. 2015). det moræneler, som dækker landskaberne i den sydøstlige del af danmark, bliver traditionelt henført til den bundmoræne, som blev afsat af det ungbaltiske isfremstød. denne aflejring af moræneler har i det sydlige og centrale østjylland navnet østjylland till formation (east jylland till forma­ tion, houmark­nielsen 1987), mens aflejringer fra det yngre genfremstød, der afsluttede isdækket over dan­ mark, benævnes bælthav till formation. typelokalitet og referenceprofil. typelokaliteten for lol­ land till formationen er grusgraven ved birket på lolland (pedersen et al. 2015). på møn findes et referenceprofil ved kobbelgård (fig. 18; houmark­nielsen 1994). lolland till formationen udgør sammen med ny borre formationen de dominerende litologier på møns landområde, nemlig omkring 80 % af dæklagene. for­ mationen kan være vanskeligt tilgængelig på grund af sin beliggenhed øverst i de relativt høje kystprofiler. i kystklinterne draperer formationen alle andre aflej­ ringer og ligger i en række af kystprofilerne inkonformt over de andre glacialtektonisk forstyrrede aflejringer. litologi. formationen består af siltet, brunt til gulbrunt moræneler (fig. 19). leret kan være massivt/kompakt, men også optræde med en svagt udviklet horisontal lag­ deling. klastindholdet er forholdsvis lavt, men der findes dog en del kalkklaster (houmark­nielsen 1994; smed 2010; gravesen et al. 2017). lag af moræneler med meget højt indhold af kalk kan klassificeres som glaci­ tektonit (banham 1977, pedersen 1988), som beskrevet i pedersen et al. (2018). fingrusanalyser fra en række lokaliteter (hyde 1986; houmark­nielsen 1994) viser, at palæozoiske kalkkorn almindeligvis dominerer over kretasiske kalkkorn. jordartssymboler: ml, ms. grænser og tykkelse. den nedre grænse er den glacial­ tektoniske inkonformitet ved formationens sål. den hviler for det meste på ny borre formationen, men i klintholm­området på kobbelgård formationen (fig. 8). den øvre grænse er på store dele af møn også landskabets overflade, men formationen er f.eks. ved tøvelde klint og hjelm nakke overlejret af sen weichsel­ ferskvandsaflejringer og stedvis af holocæne marine aflejringer, i begge tilfælde med skarpe grænser. generelt varierer formationens tykkelse fra 1 til 3 m, men boredata viser, at den kan være tykkere, selv om gla­ cialtektoniske forstyrrelser kan spille en rolle. i klinterne er målt følgende tykkelser: madses klint: 2 m, hvideklint: 2 m, tøvelde klint: 2,8 m, klintholm: 0,8 m, nordlige møns klint: 0,8 m, stubberup have: 2 m, liselund: 1,5 m, pom­ lerende; 0–2 m, hellehavn nakke: 2 m og brunshoved: mindst 2 m (hyde 1986; gravesen et al. 2017). tynde lag af formationen er også fundet på den hatformede bakke præstebjerg på det sydvestlige møn (berthelsen 1980). formationen findes ikke på de højeste dele af møns klint. udbredelse og korrelation. lolland till formation er udbredt over hele lolland, falster og møn samt på øerne i smålandsfarvandet. alder. lolland till formation har en alder på 18 000 til 17 000 år før nu (noe­nygaard et al. 2017). aflejringsmiljø. under nordøstisens tilbagesmeltning blev møn, falster, lolland og smålandsfarvandet efterladt som et småbakket morænelandskab. dette moræneplateau blev efterfølgende overskredet af den ungbaltiske is, der fra den nordøstlige del af østersøen spredte sig mod vest. analyser af klast fabrics og glacialtektonisk defor­ mation samt et højt indhold af palæozoisk kalk viser, at moræneleret er aflejret af en gletsjer, som har bevæget sig fra sydøst mod nordvest (hyde 1986). http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 22 of 53 www.geusbul let in.org med det ungbaltiske isfremstød skete der under frem rykningen en spredning af den frontale del af ismassen, som resulterede i en isbevægelse ud mod dens sider, dvs. isfremstød mod den centrale del af dan­ mark på nordsiden og mod nordtyskland på syd siden. på lolland medførte dette, at den isstrøm, som først var rettet fra øst mod vest, drejede med uret og skred frem mod nordvest. i det afsluttende stadie drej ede isbevæ­ gelsens retning endda op mod nordnordvest (pedersen et al. 2015). isfremstødet mod nordvest og nordnordvest ophørte ved den østjyske israndslinje i den centrale del af djursland, hvor gletsjerens afsmeltning gav ophav til den ekstramarginale tirstrup hedeslette (pedersen & petersen 1997). fra denne stationære linje smeltede isen tilbage med oscillerende bevægelser. oscillationen med­ førte bl.a. bælthav isfremstødene, hvor en serie lokale israndslinjer er kortlagt (milthers 1948). på møn gav dette fremstød kun ophav til en aflejring af moræneler, mens den på dele af lolland og falster gav ophav til opsplitningen af lolland till formationen i to enheder af moræneler adskilt af et tyndt lag smeltevandssand. sen weichsel (senglacial, mis 2) aflejringer fra den sidste del af sen weichsel (senglacial) kendes især fra to lokaliteter på møns sydkyst: hjelm og tøvelde, hvor ferskvandslagene omfatter store dele af det senglaciale tidsafsnit: ældste dryas, bølling, ældre dryas, allerød og yngre dryas (kolstrup 1982; noe­ nygaard & heiberg 2001). aflejringerne er dannet mel­ lem 17 000 og 11 700 år før nu. tøvelde er en af de bedst undersøgte senglaciale lokaliteter med et næsten helt komplet senglacialt profil (fig. 20). senglaciale ferskvandslag er også fundet i moser på høje møn (hintze 1937), mens der ikke er tegn på marine aflejringer af denne alder på møn. ved madses klint findes lag af kalkgytje og tørv på toppen af den ungbaltiske lol­ land till formation. lagene er ikke dateret, men kan være senglaciale eller postglaciale (fig. 21). endelig træffes senglaciale aflejringer i bunden af borre sømose pro­ filet (mikkelsen 1949). ferskvandssedimenterne består overvejende af ferskvandsler (tl), ferskvandssilt (ti), fersk vandsgytje (tp), ferskvandstørv (tt), vekslende lag (tv), ferskvandssand (ts) og ferskvandsgrus (tg). lolland till fm fig. 18 gråbrunt moræneler fra lolland till formationen i klinten ved kraneled, klintholm. fig. 18. greyish brown clayey till of the lolland till formation in the uppermost part of the cliff at kraneled, klintholm. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 23 of 53 www.geusbul let in.org tøvelde­lokaliteten er velblottet over ca. 100 m i kystklinten ved tøveldestenen og er sammensat af flere små søbassiner, som har udviklet sig til en sø gennem tid. hjelm­lokaliteten har en mere begrænset udstrækning, og den ligger højt i kystprofilet og er van­ skeligt tilgængelig (heiberg 1991). de to velbeskrevne lokaliteter viser søbassiner med en omfattende fersk­ vandsfauna og ­flora samt også terrestriske dyr, der har levet omkring søerne (johansen 1904; heiberg og bennike 1997; noe­nygaard & heiberg 2001). faunaen omfatter bl.a. terrestriske pattedyr (gnavere), padder, snegle, muslinger, insekter og planter, samt ferskvands­ fisk, snegle, muslinger, insekter og planter. den samlede tykkelse er på op til 3,8 m. på høje møn er aflejringer i flere moser blevet undersøgt (hintze 1926, k. jessen i hintze 1937), og i nogle af dem er der fundet en sen weichsel­ferskvands­ lagfølge af ler fra ældre dryas, gytje fra allerød og ler fra yngre dryas, overlejret af lag fra holocæn. udviklingen i sen weichsel (senglacial) i sen weichsel blev en stor isdæmmet smelte vandssø, den baltiske issø, etableret i den sydlige del af østersøen for omkring 16  000 år siden. i de mindre sølavninger på møn ved hjelm og tøvelde blev der med lidt varier­ ende udvikling dannet forskellige sedimenter, især tørv og gytje. søbassinerne blev dannet i et højtliggende landskabsparti og opstod ved smeltning af dødis efter­ ladt fra afsmeltningen af den ungbaltiske is. de klima­ tiske forhold havde stor betydning for aflejringen i de relativt lavvandede søer, men også niveauændringer i den baltiske issø spillede en rolle for søernes udvikling (kolstrup 1982; noe­nygaard & heiberg 2001; noe­ nygard et al. 2017; rosentau et al. 2017). i ældste dryas indvandrede en lav tundravege­ tation til et landskab uden planter. i de fladbundede lavninger ved hjelm og tøvelde blev der aflejret sand, silt og resedimenteret moræneler med plantemateriale, som blev skyllet ud eller gled ud som flydejord fra det omgivende landskab. plantematerialet bestod bl.a. af dværgpil, bynke, græsser og halvgræsser. i bølling tid blev det varmere, og planter som havtorn fig. 19 gråbrunt moræneler tilhørende lolland till formationen under sen weichsel lakustrine aflejringer ved tøvelde klint. ved toppen af de lakustrine sedimenter ses en kalkgytje. fig. 19. grey-brown, clayey till of the lolland till formation below late weichselian freshwater deposits at tøvelde klint. at the top of the lacustrine succession a calcareous gyttja is seen. fig. 20 sen weichsel – holocæne aflejringer i et søbassin ved tøvelde, som ligger i en lavning i moræneler fra det ungbaltiske isfremstød. fig. 20 late weichselian – holocene lacustrine deposits at tøvelde, formed in a shallow depression of the till of the young baltic ice advance. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 24 of 53 www.geusbul let in.org og dværgbirk indvandrede sammen med pil, dryas og padderokke. det højere planteindhold betød et højere humusindhold i jordbunden, og øget nedbør førte til udvaskning af kalk fra det kalkholdige moræneler og ­sand ud i de lavvandede søer, som var etableret. fiske­ fossiler er truffet i søerne, og landsnegle levede omkring dem. humusholdigt sand blev aflejret i et sumpet sømiljø i den sidste del af bølling. i ældre dryas blev klimaet mere tørt og koldt, og mængden af dværgbirk faldt markant. kalkholdigt ler med dropsten blev aflejret i søbassinerne, og der blev skyllet usorteret jordmateriale ud i bassinerne, hvor kalkudvaskningen stoppede. i allerød blev det betydeligt varmere end i de tidligere tidsafsnit, selv om to korte kuldeperioder indtraf. almin­ delig birk indvandrede. gytjelag og tørv blev dannet i de fladvandede søer ved tidsafsnittets begyndelse. udvask­ ningen fra de kalkholdige lag af moræneler blev forøget, og det højere kalkindhold i søvandet bevirkede, at de lakustrine sedimenter blev præget af kalkgytje overvej­ ende bestående af kransnålalger (characékalk). disse alger kan optage ca2+­ioner fra søvandet og udfælde søkalk. vandspejlet i søerne varierede gennem tidsafsnit­ tet, men mod slutningen faldt det med sumpdannelse til følge, og der blev igen aflejret alge­ og detritusgytje (fig. 20). i og omkring søerne fandtes en artsrig flora og fauna. i det sidste sen weichsel­tidsafsnit, yngre dryas, blev det igen koldere, og aflejringerne blev mere sandede med laminerede silt­ og lerlag uden characékalk, men med dropsten og talrige planterester. søernes vand­ niveau var stigende gennem tidsafsnittet. sedimentatio­ nen markerer opfyldningen af søerne med materiale fra det omgivende træløse landskab. ved år 11 700 før nu blev sen weichsel afløst af det generelt varmere klima i holocæn. under den baltiske issøs forskellige faser fra weich­ sel til holocæn gennemgik farvandene omkring møn, fakse bugt og hjelm bugt markante niveauændringer. der blev dannet lagune­, ferskvands­ og kystaflejringer samt sø­ og moseaflejringer gennem sen weichsel (jen­ sen 1993; bennike & jensen 1995; jensen et al. 1997). holocæne aflejringer (postglacial, mis 1) det holocæne tidsafsnit (postglacial) siden 11 700 år før nu er opdelt i tidsafsnittene præboreal, boreal, atlan­ tisk, subboreal og subatlantisk tid (nutid). disse enheder er bestemt på få lokaliteter, hvor profilerne ofte består af både ferskvands­ og marine aflejringer. det mest fig. 21 sen weichsel ferskvandsaflej­ ringer højt oppe i klinten ved madses klint, som indeholder en ca. 20 cm tyk kalkgytje. foto: tove stockmarr. fig. 21: the late weichselian freshwater deposits high up in the section at madses klint, which contains a bed of 20 cm thick calcareous gyttja. photo: tove stockmarr. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 25 of 53 www.geusbul let in.org komplette profil findes i borre sømose lavningen (fig. 22; mikkelsen 1949), mens dele af de holocæne aflej­ ringer er undersøgt i forbindelse med de senglaciale profiler ved tøvelde (noe­nyegaard & heiberg 2001). ferskvandsaflejringer de holocæne ferskvandsaflejringer består af fem forskel­ lige sedimenttyper: tørv (ft), gytje (fp), ler (fl), sand (fs) og heterolitiske (fv) aflejringer. i alt dækker ferskvand­ saflejringerne ca. 7 % af landarealet. ferskvandsaflejrin­ gerne af holocæn alder består af to hovedgrupper: 1) på toppen af og mellem de marine postglaciale aflejringer og 2) i de isolerede lavninger (moser og søer) i det kupe­ rede morænelandskab, ofte på senglaciale lag. på fig. 22 ses, hvordan de postglaciale lag overlej­ rer de senglaciale aflejringer i borre sømose, hvor en næsten kontinuerlig lagfølge er aflejret fra præboreal til subatlantikum. det geologiske kort viser et stort antal isolerede forekomster af tørv, der er dannet i de død ishuller, som opstod ved slutningen af sengla­ cialtiden, da resten af den sidste smeltende gletsjer forsvandt. andre spredte ferskvandsaflejringer opstod i flade lavninger på morænefladen og i det småbakkede morænelandskab. større søer som røddinge, råby og råbylille søer groede til i holocæn, og der blev dannet tørv som i busemarke mose. det største ferskvandsbas­ sin er borre sømose, som behandles nærmere i afsnittet om den holocæne udvikling. på høje møn er der i moser over de senglaciale lag fundet aflejringer fra holocæn startende med boreal (hintze 1926, jessen i hintze1937). ferskvandsaflejringerne domineres som nævnt af tørv. undersøgelser af tørven viser en vegetations­ udvikling gennem præboreal, boreal, atlantisk, sub­ boreal og subatlantisk tid, som svarer til den øvrige del af det centrale og sydlige danmark (noe­nygaard et al. 2017). store dele af et bjørneskelet er fundet i kam mose. bjørnen formodes at have levet i den tidlige del af holocæn (jessen 1929). marine aflejringer de holocæne marine aflejringer udgør ca. 10 % af jordarterne på land og opdeles i gytje og organisk mud­ der (hp), ler (hl), sand (hs) samt grus og sten i strand­ volde (hg). havaflejringer træffes mange steder langs kysterne med strandvolde som den dominerende del. strandvoldene danner vinkelforland og krumodde sy­ stemer på øens fremskudte kyster og pynter, som ved f.eks. hårbøllebro. ulvshale er en stor strandvoldsslette, hvor der under de grovkornede strandvolde ligger omkring 7,5 m marint sand med postglaciale skaller og højt indhold af organisk materiale. nyord er for­ modentlig dannet som et tilgroningsforland beskyttet af den tilknyttede knold af 20 m tykt moræneler (fig. 23). knolden er antagelig dannet af glacialtektonisk opskudte flager af moræneler. på mange grovkornede forstrande og strandvolde dominerer flint, som stam­ mer fra skrivekridtklinterne. på stranden for foden af møns klint kan træffes flere generationer af strand­ volde (fig. 24). strandvolde kendes også fra nordkysten af møn og ud mod gyldenløves flak (andersen 1936). tykke havaflejringer er dannet under havstigningerne i atlantisk tid (littorinahavet). de er især fundet i borre lavningen i form af cardiumgytje (efter muslingen cardium edule; mikkelsen 1949). den hævede havbund fra littorinahavet, som ligger i kote 0 eller –1, er afgrænset fra det nuværende hav med diger og kan udgøre store, næsten vandrette flader, som ved kostervig, hvor fladen er opdyrket. 0 -5 -10 2 0 100 200 300 400 500 600 m 1 3 4 5 6 7 8 9 1110 12 13 14 15 16 17 18 19 20 21 vest øst d yb de (m ) subboreal-subatlantisk vekslende gytjelag atlantisk cardiumgytje præboreal ? lersubatlantisk tørv boreal-atlantisk gytje og tørv weichsel moræneler fig. 22 fig. 22. profil gennem borre sømose, der viser udviklingen fra den allersidste del af sen weichsel til den sidste del af holocæn. tallene 1–21 angiver placeringer af korte boringer. fra gravesen et al. (2017), modificeret fra mikkelsen (1949). fig. 22. section through borre sømose showing the development from the late late weichselian towards the last part of the holocene. the numbers 1–21 show the positions of shallow boreholes. from gravesen et al. (2017), modified from mikkelsen (1949). http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 26 of 53 www.geusbul let in.org klitdannelser nutidige flyvesandsdannelser (es) i form af kystklit­ ter findes på nordøstkysten af ulvshale, langs kysten ved klintholm havn og ved råbylille i hjelm bugt. der findes også ganske få områder med flyvesand oven på ferskvandsaflejringer. udviklingen gennem holocæn i præboreal i den første del af holocæn steg tempera­ turen hurtigt. herefter fulgte en kort kuldeperiode, inden det igen blev varmere og tørt. græs og urter vandrede ind. i den fladvandede sø ved tøvelde blev der først aflejret algegytje med højt organisk indhold, og efter en stigning af vandspejlet blev der aflejret characé kalkgytje (mikkelsen 1949; noe­nygaard & heiberg 2001). i den kolde periode blev ler skyllet ud over gytje lagene, og nye gytjelag blev aflejret igen, da det blev varmere. nu ind­ vandrede bl.a. birk, fyr og asp og senere bøg, eg, elm og hassel. i borre sø lavningen, der blev dannet som en u­formet smeltevandskanal ved smeltningen af den ungbaltiske is, fandtes en fladvandet sø ved overgangen til holocæn (fig. 22). i præboreal og boreal tid blev der dannet en mose med trævækst. i præboreal var danmark landfast med sydsverige og sydengland. denne situation fortsatte et stykke ind i boreal tid og kaldes samlet fastlandstiden. den baltiske issø eksisterede stadig ved starten af præboreal, men den blev tappet 10 300 år før nu. i boreal tid voksede planter ud over tøvelde søerne, eller træer væltede ned i dem, og der blev dannet tørv, heraf en del tykke lag af sumptørv med tusinder af skaller af ferskvandssnegle og ­muslinger. søerne blev efterfølgende omdannet til lavmoser. i dette tidsafsnit blev det baltiske område igen til en stor ferskvandssø, ancylussøen, der mindede om den baltiske issø, der havde eksisteret få tusind år forinden. derefter, i atlantisk tid, blev store dele af danmark overskyllet af littorinahavet. på møn blev lavtliggende områder langs kysten og gamle fjorde som borre sømose en del af dette hav (fig. 22; krog 1979). der blev aflejret tykke marine lag af cardiumgytje, og aflejringen fortsatte op gennem atlantisk og videre ind i subboreal tid. fra disse tidsafsnit kendes en omfattende flora og fauna. i subatlantisk tid var der afvekslende lakustrine for­ hold og marine fjorde, og aflejringerne ændredes i takt hermed mellem ferskvandsgytje og kalkgytje. de skiftende aflejringsforhold var også et resultat af til­ stedeværelsen af tærskler og barrierer, som stedvist blev dannet foran fjordene som f.eks. i borre sømose i takt med, at havniveauet ændrede sig. til sidst blev der mange steder dannet søer, som voksede til, dannede tørv og blev til moser. de recente marine aflejringer består mest af sand, grus og sten. ikke mindst strandvoldene er domineret af grus og sten. havbunden og de underliggende lag omkring møn er kortlagt i forbindelse med råstof­ undersøgelser efter sand, grus og stenressourcer (fred­ ningsstyrelsen 1977, 1986) i henholdsvis hjelm bugt og fakse bugt og sammenstillet af jensen (1993). i farvandet mellem falster og møn findes grønsunds tragtformede munding med en central rende omgivet af flak og bar­ rer dannet af stormskabte strømme (fredningsstyrelsen 1977). den yngste nutidige aflejring er flyvesand, som har dannet klitter, især langs kysterne. lagenes optræden og strukturelle forhold skrivekridtets beliggenhed og fordelingen af skrive­ kridtlag overlejret af kvartære enheder, samt opbyg­ ningen af den kvartære lagserie og den strukturelle fordeling af de forskellige formationer er dokumen teret i fire tværsnit a–b, c–d, e–f og g–h gennem landskabet fig. 23 holocæn strandvoldsslette ved nyord–ulfshale set mod øst. fig. 23. holocene plain of beach ridges at nyord–ulfshale. view towards east. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 27 of 53 www.geusbul let in.org på møn. disse indgår på kortbladet møn. desuden er to kystprofiler, henholdsvis møns klint (hi storisk) og hvide klint, medtaget i beskrivelsen. konstruk­ tionen af tværsnittene er baseret på boringer inden for kortbladet, som findes i geus’ boredatabase jupi­ ter (gravesen & fredericia 1984) samt oplysninger fra grundvandskortlægning (rambøll 2006, 2007), råstof­ grave (region sjælland 2012) og klintprofilerne. den vestlige del af tværsnit g–h er vist her som et særskilt tværsnit med de nye formelle stratigrafiske enheder anført på de enkelte lag (fig. 25). på profilerne fra kyst­ blotningerne er der foretaget en strukturel tolkning af lagseriernes forløb i dybet. tværsnittene baseret på boringer repræsenterer et overblik over undergrundens opbygning fra toppen af skrivekridtet til terrænoverfladen på den indre del af møn (fig. 26). møns klint vises i fig. 27 i en bearbejdet udgave af det klassiske profil udarbejdet af puggard (1851). profilet langs hvideklint er baseret på en ny opmåling understøttet af opmålingen af stockmarr (1996). det er dog kun de østligste 800 m af kystklinten fra hvideklint til hjelm nakke, som indgår i profilet (fig. 28). a b fig. 24 generationer af holocæne strandvolde ved gråryg, møns klint. a: overblik over lokaliteten med recente strandaflejringer i forgrunden. b: detal­ jer af de ældste strandvolde. fig. 24. generations of holocene beach ridges at gråryg close to møns klint. a: overview of the locality with recent beach deposits. b: details of the oldest beach ridges. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 28 of 53 www.geusbul let in.org de geologiske tværsnit baseret på boringer og andre kortlægninger tværsnit a–b tværsnit a–b (orienteret nv–sø) går på tværs af det vestlige møn, fra hvideklint mod nordvest over nørre frenderup til koster vig, over koster halvøen til ulvsund og videre til gammel kalvehave på sjælland. det småbakkede morænelandskab ved gammel kalvehave og koster land er omkranset af et holocænt marint forland. ved damsholte krydser tværsnit a–b en dal, som løber sydvest–nordøst. dalen er fyldt med smeltevandsler, ­sand og ­grus afsat af hovedfrem­ stødet. den tolkes som en tunneldal (rasmussen 1965; smed 2014). dette dalforløb kan følges mod nordøst ud gennem stege nor, men som det fremgår af tværsnit a–b, er dalens forløb delvis modificeret af det baltiske isfremstød. stege nor og bakkerne i landskabet sydøst herfor danner et sammensat parallelt dal­ og bakkestrøg, som skyldes isfremstødet fra sydøst af den ungbaltiske is. i den sydøstligste del af tværsnit a–b findes gen­ deformerede, opskudte skiver af skrivekridt i hvideklint. skrivekridtoverfladen, som ikke er påvirket af glacialtek­ tonisk oppresning, ligger nogenlunde horisontalt mellem kote –20 m og –30 m (fig. 2), hvilket i tværsnit a–b dog kun er tilfældet ved æbelnæs sydøst for kostervig. tværsnit c–d tværsnittet c–d, som er orienteret n–s, starter i hjelm bugt ved råbylille sø og går nordpå tværs over den øst­ lige del af møn, op over højdedraget ved elmelunde og videre til nordkysten ved nordfelt vest for borre sømose. det bakkede morænelandskab udgøres af lolland till formationen fra det ungbaltiske isfrem­ stød, men omkring det højtliggende elmelunde, som er presset op af glacialtektoniske skrivekridtskiver, træder smeltevandssand og grus fra ny borre formationen frem i terræn. de to formationer kan følges gennem hele profilet. mod syd træffes holocæne ferskvands­ aflejringer i den ø–v­orienterede råbylille sø. den ligger i en fordybning i morænelandskabet, der kan tolkes som en del af råbylille tunneldalen. søen er groet til, og der er dannet tørv. langs kysten træffes nutidigt flyvesand. den upåvirkede skrivekridtoverflade ligger nogenlunde horisontalt i kote –30 til –40 m. tværsnit e–f tværsnittet e–f, orienteret sv–nø, går fra stege over damsholte, røddinge sø og store damme til fanefjord. steges morænelandskab består af mindre aflange bak­ ker orienteret sv–nø, som ligger på skiver af opskudt skrivekridt. i hele profilets længde følges en dalstruktur, som har fungeret som dræningskanal for nø­isens smelte­ vand. dele af dalen er tolket som en tunneldal med mindre åse, bl.a. fanefjord ås og hatformede bakker (rasmussen 1965; smed 2014). hele systemet har senere fungeret som dræningskanal for den ungbaltiske is bl.a. gennem rød­ dinge sø kanalen, og ny borre formationen blev aflejret. røddinge lavningen er nu tilgroet, og der er dannet tørv. fanefjord omkranses af holocænt marint sand, og ved østersøkysten findes grovkornede strandvolde. ved en kabelnedgravning nær transformatorstationen i den østlige udkant af neble blev skiver af skrivekridt blotlagt i sommeren 2015 (fig. 26). det viser, at de skrivekridtskiver, der er antydet i den østlige tredje del af tværsnit e–f, nok er underdrevet i profilkonstruktionen, som er udført 10 år tidligere. som det ses i fig. 26, når skrivekridtet helt op til overfladen. vest for landsbyen neble er skrivekridtoverfladen så godt som upåvirket af glacialtektonik. den prækvartære overflade ligger nogen­ lunde horisontalt i en dybde omkring kote –25 til –30 m. ny borre fmny borre fm lolland till fm ristinge klint till fm kraneled fmmidtdanske fmmidtdanske till fm kobbelgård og klintholm fm lolland till fm ny borre fm klintholm till fm lolland till fm midtdanske till fm ristinge klint till fm midtdanske till fm ny borre fm lolland till fm kobbelgård fm skrivekridt fig. 25 geologisk tværprofil g–h med formationsnavne for de kvartære aflejringer (sammenlign med tilsvarende profil og signatur­ forklaring på kortbladet). dette og de øvrige profiler vist på kortbladet er konstrueret og tolket ved hjælp af boredata fra kortblads­ området. de kvartære aflejringer domineres af smeltevandssand og grus (rød), smeltevandsler og silt (orange) og moræneler (brun). tværprofilet løber øst–vest gennem midten af møn mellem stege og høje møn. fig. 25. geological profile g–h, annotated with formation names and other geological units (compare with the corresponding profiles and legend on the map sheet). this and the other profiles shown on the map sheet have been constructed and interpreted based on well data from the map sheet area. meltwater sand and gravel (red), meltwater clay and silt (orange) and clayey tills (brown) dominate the quaternary deposits. the profile is orientated e–w through the central part of møn from stege to høje møn. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 29 of 53 www.geusbul let in.org tværsnit g–h på langs af den centrale del af møn findes det længste tværsnit g–h (orienteret v–ø), som løber fra stege over råbylille, råbymagle, borre sømose, busemarke, store klinteskov til møns klint ved store klint og dronninge­ stolen. et udsnit af dette tværprofil er vist på fig. 25, hvor høje møn er udeladt. de litostratigrafiske enheder, som er opstillet i denne publikation, er markeret på fig. 25. mod øst i høje møn­området er de store skrive­ kridtskiver presset op til 125 m o.h., og der ligger et tyndt dække af glaciale, interglaciale, senglaciale og postgla­ ciale aflejringer over skiverne. det høje terræn afløses af den markante fordybning med borre sømose, som er fyldt op af senglaciale og postglaciale aflejringer. borre sømose fordybningen er desuden fyldt med afl ej ringer af sand og grus tilhørende ny borre formationen og har fungeret som dræningskanal for den ungbaltiske is. området syd for elmelunde er et højere beliggende småbakket morænelandskab. i stege nor omkranser holocæne marine sandaflejringer havbugten. på en del af profilet i fig. 25 ses, hvordan de lito­ stratigrafiske enheder antages at ligge i forhold til hinanden inde på møn, altså væk fra kystprofilerne, som de ikke direkte kan sammenlignes med på grund af det forskellige datagrundlag. tværprofilet viser føl­ gende formationer: ristinge till, kraneled, klintholm till, kobbelgård, midtdanske till, ny borre og lolland till formationerne. vest for høje møn ligger det meste af skrivekridtoverfladen nogenlunde horisontalt i kote –30 til –40 m, men kan stedvis nå op til ­20 m. under høje møns skrivekridtskiver mod øst forventes den faststående skrivekridtoverflade at ligge under kote –100 m. isolerede skrivekridtlegemer af forskellige størrelser er indlejret i de kvartære aflejringer (se også hou mark­nielsen 2003). skrivekridtprofiler i kystklinter møns klint møns klint er et imponerende kystprofil, som giver et nord–sydgående erosionssnit gennem det østlige møns stærkt kuperede og efter danske forhold høje, stejle landskab, det såkaldte høje møn. i mere end 150 år har det været velkendt, at klintens morfologi skyldes dens opbygning af skrivekridtflager, som er stablet oven på hinanden (puggaard 1851; schou 1949; pedersen 2000). den strukturelle dannelse blev allerede for mere end fig. 26 højtliggende skrivekridtskiver blottet i en udgravning ved transformatorstationen i neble. fig. 26. elevated chalk sheets exposed in an excavation at the transformer station in neble. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 30 of 53 www.geusbul let in.org fi g. 2 7 m øn s kl in t­ pr ofi le t, so m d et b le v op m ål t a f c . p ug ga ar d om kr in g 18 50 . fi g. 2 7. t he m øn s kl in t p ro fil e as m ea su re d by c . p ug ga ar d at a ro un d 18 50 . http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 31 of 53 www.geusbul let in.org 100 år siden tolket som resultatet af glacialtektonik. fra den skandinaviske iskappe rykkede i sidste istid en isstrøm frem gennem østersøen mod kattegat. ved isens front blev kridtskiverne med overliggende kvartære aflejringer skubbet op i glacialtektoniske kom­ plekser på møn og rügen (johnstrup 1874; gripp 1948). siden da har møns klint været betragtet som en nøgle­ lokalitet for glacialtektonik (fig. 27; slater 1927; aber et al. 1989; pedersen 2000, 2014). i en overordnet strukturel sammenhæng kan møns klint inddeles i et imbrikeret proksimalt kompleks mod syd, en antiform stak i den centrale del og lavt hældende forlandsoverskydninger i den distale del mod nord (peder­ sen 2000). desuden er især den nordlige del af klinten påvirket af overpræget deformation, som skyldes det sid­ ste afsluttende isfremstød fra skåne, inden isen smeltede fuldstændig væk fra danmark for ca. 16 000 år siden. det imbrikerede kompleks i den proksimale, sydlige del af klinten er karakteriseret af enkelte stejltstående skiver som f.eks. hvidskud, st. stejlbjerg og sommerspir. de struk­ turelle forhold for disse skiver er vist i blok diagrammet på det geologiske kortblad (vignetten ’stratigra fisk og struk­ turel opbygning af møns klint’). det skal her bemærkes, at décollement­niveauet ligger ca. 100 m under havover­ fladen (fig. 42; pedersen 2000). det allersydligste parti er jættebrinken, som består af tre skrivekridtskiver, der er skudt op over hinanden langs lavtliggende overskyd­ ninger og nu har en fladtliggende orientering (fig. 27). syd for jættebrinken forekommer glaciale sedimenter i et hul, hvor den underliggende skrivekridtoverflade er under havniveau. stien til stranden fra fyret på den sydlige pynt af møns klint går ned gennem den glaciale lagserie, som dog er skjult af tæt krat. klintens sydlige del tolkes som det sted, hvor der var kontakt mellem den fremadskridende is og den deformerede lagpakke. det moræneler, som er aflejret her, tolkes som tilhørende lolland till formation, som ved en længerevarende aflejring af bundmoræne i den tektoniske depression bag de opskudte skiver har fået en forøget tykkelse. hundefangsfaldet adskiller jættebrinken fra hunde­ fangsklinten længere mod nord, der er den første af de stejltstående klinter i den proksimale zone (fig. 27, 42). denne zone strækker sig frem til sommerspir pynten, hvor den brede rende sandfaldet har fået sit navn efter den store mængde smeltevandssand, som findes her. sandfaldet tolkes som et ’piggyback’ bassin, hvor en smeltevandsflod gennemstrømmede dalen mellem sommerspirskiven og skiven nord for sandfaldet, mens skiverne blev bevæget fremefter af det glacialtektoniske pres. skiven nord for sandfaldet strækker sig fra græ­ deren og frem til maglevandsfaldet (fig. 27). langs dens sider pålejrer sandet undersiden af sommerspirskiven og oversiden af skiven ved græderen. sandet er sam­ tidig moderat foldet i en synklinal som et slæb langs overskydningen, hvilket viser, at umiddelbart efter san­ det blev aflejret, blev det deformeret af de igangværende glacialtektoniske processer. dronningestolens strukturer tolkes som et resul­ tat af en overskydning af flere skiver, som i deforma­ tionens tidlige fase blev placeret fladt oven på hinanden (fig. 27 & 42). efterfølgende blev alle skiverne skubbet med op over en dybtliggende rampe, hvorved de blev bøjet i en hængende­blok antiklinal struktur (pedersen 2000). disse forhold fremgår af en detaljeret strukturel tektonisk brecciering overskydning lolland till formation midtdanske till formation møns klint formation klintholm till formation kraneled formation 0 20 m -20 -40 -60 -80 hk 11 hk 13 hk 12 hk 10 hk 9 hk 8 hk 7 hk 6 hk 5 hk4 hk3 hk2 hk1 sv nø 0 20 m -20 -40 -60 -80 100 m 0 20 m -20 -40 -60 -80 hk11 hk13 hk12 hk10 hk9 hk8 hk7 hk6 hk5 hk 4 hk 3 hk 2 hk 1 sv nø 0 20 m -20 -40 -60 -80 40 80 1000 20 hk5 hk4 stubberup have formation kobbelgård formation fig. 28 strukturen af hvideklint baseret på nye opmålinger suppleret med informationer fra stockmarr (1996). hk1–hk12 refererer til skrivekridtskiver nummereret fra nordøst mod sydvest. disse betegnelser er også brugt i fig. 29–39 fra hvideklint. fig. 28. the structure of hvideklint based on new measurements supplemented by information from stockmarr (1996). hk1 to hk12 refer to individual chalk sheets numbered from northeast to southwest. these are also used in figs 29–39 from hvideklint. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 32 of 53 www.geusbul let in.org undersøgelse neden for geocenter møns klint (peder­ sen 2003; pedersen & gravesen 2009). i den distale del af det glacialtektoniske kompleks har overskydningerne en moderat hældning mod syd (fig. 27, 42). overskydningerne går ind over forlandet, som består af ler tilhørende kobbelgård formationen; det fremgår af en opmåling ved nonnebrinken–liselund foretaget under forfatternes feltarbejde. en særlig struk­ tur er her en ’roll­over’ antiklinal, som er blottet neden for nonnebrinken. hvideklint hvideklint er en mere end 1 km lang kystklint ved den vestlige del af hjelm bugt. klintprofilet stryger nø–sv og giver et erosionssnit gennem bakkelandet på den syd­ vestlige del af møn (fig. 28). kystklinten er omkring 20 m høj og består af mere end 13 skrivekridtskiver med mellemlejrede glaciale aflejringer, der på lang afstand ses som en lysende hvid klintstrækning. opbygningen af hvideklint og dens interne struk­ turers arkitektur er betinget af tre glaciodynamiske forhold: (1) skiver af skrivekridt blev først skudt op af hovedfremstødet fra nordøst mod sydvest. (2) kom­ plekset af opskudte skrivekridtskiver blev afslutningsvis deformeret af det ungbaltiske isfremstød fra sydøst mod nordvest. (3) kombinationen af skrivekridt og ler medførte en intensiv brecciering og dannelse af mudder­ diapirer (aber 1979; berthelsen 1979; stockmarr 1996). et a b fig. 29 folder i hvideklint. a: over­ kippet synklinal i skive hk3 med ø–v­orien teret foldeakse. lyse glacio­ fluviale sandlag tilhørende stubberup have formationen er indfoldet i det mørkere kobbelgård formation ler. b: næsen af den overkippede synklinal med vekslende lag af ler, silt og sand. fig. 29. folds at hvideklint. a: overturned syncline in sheet hk3 with an e–w-oriented fold axis. the light glaciofluvial sand of the stubberup have formation is folded into the darker clay of the kobbelgård formation. b: the nose of the overturned syncline includes alternating layers of clay, silt and sand. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 33 of 53 www.geusbul let in.org tværsnit af hvide klints strukturer vist i fig. 28 er baseret på opmåling af klinten, målinger af stryg ning og hæld­ ning for lagplaner og overskydningsplaner, måling og konstruktion af foldeakser kombineret med overvejelser relateret til skrivekridtskivernes volumen (balancerede tværprofiler, pedersen 2005). denne tolkning støttes af det konstruerede kort over skrivekridt overfladens højde­ forhold (fig. 2) og tværsnit e–f på kortbladet, som begge er baseret på boredata fra især det vestlige møn. det tolkede tværsnit kan tillige opfattes som en model for bakkelandet omkring hjelm og nørre frenderup. tværsnittets nulpunkt er ved en bæk, som løber ud over stranden neden for ørbæklund. de første (nordøstlige) ca. 300 m af hvideklint (hk) domineres af skrivekridtskiverne hk1 og hk2. på ryggen af hk1 fore­ kommer de nedre kvartære enheder af kobbelgård for­ mationen i en smal, overkippet synklinal. denne foldning har tillige påvirket den øverste del af skrivekridtet i hk1, som dog ved forskydningsspænding er omdannet til en kridt­glacitektonit. foldningen afskæres af en overskyd­ ningsflade, hvorpå den midtdanske till formation hviler. på toppen overlejrer lolland till formationens bund­ moræne hele kystprofilet med en diskordant grænse. overskydningsgrænsen mellem hk1 og hk2 er ikke altid tydeligt blottet, men kan identificeres ved, at den syd­ vestlige flanke i hk1s hængende antiklinal står næsten ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ hk7 skrivekridt skrivekridt overskydning kraneled fm hk6 erosionsflade fig. 30 overskydning i hvideklint, som adskiller skiverne hk6 og hk 7 med en isoklinalt foldet synklinal af ler fra kraneled formationen. overskydnin­ gen kan følges gennem hele fotoet. fig. 30. thrust fault at hvideklint separating sheets hk6 and hk7 with an isoclinally folded syncline consisting of kraneled formation clay. the thrust fault can be traced throughout the photo. hk8 hk7 hk8 hk9 fig. 31 overskydninger mellem skive hk7, hk8 og hk9 i hvideklint med fronten af skive hk9. fig. 31. thrust faults between sheets hk7, hk8 and hk9 at hvideklint together with the front of sheet hk9. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 34 of 53 www.geusbul let in.org vinkelret på ryggen af en svagt nø­hældende lagstilling i overfladen af hk2. forholdene er mere vanskelige at udrede hen mod næsen af hk2, hvor glacialtektonisk brecciering bliver mere og mere fremherskende. næsen af hk2 er skudt op langs en ca. 45° nø­hældende over­ skydning, hvor kobbelgård formationens ler er over­ kippet på sydvestflanken af den opskudte hængende­blok antiklinal. mellem blotningen af skrivekridtet i næsen af hk2 og en anden blotning af skrivekridtet i ryggen af hk3 findes en 120 m lang strækning med tilskredede og tilgroede glaciale aflejringer. i den nordøstlige del af denne strækning optræder en overkippet synklinal. dens foldeakse har retningen ø–v svarende til, at opskydningen er sket fra syd (fig. 29). synklinalen og den hermed forbundne overskydning hælder stejlt mod syd. det er således en overpræget foldning, hvor struk­ turer hidrørende fra de nordøstfra opskudte skiver er redeformeret af en opskydning ved fremrykningen af den ungbaltiske is fra syd mod nord. den fra syd over­ kippede synklinal aftegner sig tydeligt i de laminerede og tyndbænkede smeltevandssedimenter i kobbelgård formationen (fig. 29). op mod ryggen af hk3 forekommer tre til fire imbrike­ rede skiver domineret af leret i kraneled, klintholm till og kobbelgård formationerne. da der ikke er blotninger hk8 fig. 32 intenst deformeret skrivekridt i skive hk8 ved basis af hvideklint, skif­ tende fra skrivekridt­glacitektonit til brec­ cierede skrivekridt­moræneaflejringer. fig. 32. intensely deformed chalk in thrust sheet hk8 at hvideklint, grading from chalk-glacitectonite to brecciated chalk-till deposits. hk10 hk9 ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ fig. 33 foldede og opskudte lag i hvideklint, der forsætter dele af krane­ led, klintholm till og kobbelgård for­ mationerne mellem skiverne hk9 og hk10. fig. 33. folded and thrust-displaced layers at hvideklint, representing parts of the kraneled, klintholm till and kobbelgård formations between the sheets hk9 and hk10. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 35 of 53 www.geusbul let in.org af skrivekridt mellem lerskiverne, er de tolket som tek­ tonisk båret på ryggen af hk3. bedømt ud fra profiltolk­ ningen har skiverne deres rod langs inkonformiteten på toppen af skrivekridtet 35 m u.h. imidlertid må det forventes, at décollement­niveauet for selve skiverne af skrivekridt ligger mindst 15 m dybere, svarende til skivernes maksimale tykkelse som de er blottet i klinten. på tværprofilet antydes det, at den basale décollement­ zone ligger i kote –80 m, men i denne dybde ville kun skrivekridt forskubbes hen over faststående skrivekridt. skrivekridtskiverne hk3, hk4 og hk5 danner et sammenhængende kompleks, der kan tolkes som en ’antiformal stack’. det vil sige, at skiverne først er skubbet lateralt hen over hinanden, inden de sammen er skudt op over den underliggende rampe under hk5. hen over næsen af hk4 optræder en splint af en skive, der tol­ kes som en opskydning af skrivekridt skubbet op fra syd. skrivekridtskiven hk6 er en tynd skive, der er skudt ind over den relativt tykke skive hk7. den overskydning, som adskiller de to skiver, optræder sammen med en isoklinalt foldet synklinal af mørkt ler, der kan henføres til kraneled formationen (fig. 30). skrivekridtet i hk6 og den øverste del af hk7 er opbrudt og breccieret og kan bedst betegnes som en skrivekridt­glacitektonit. næsen af hk7 er skudt op over ryggen af overskydningsskiven hk8, hvor flere kvartære lag er foldet i en overkippet hk10 lolland till fm midtdanske till fm klintholm till fm fig. 34 skrivekridt ved fronten af skive hk10 i hvideklint overlejret af en serie kvartære lag omfattende kraneled, klintholm till, kobbelgård, midtdanske till og lolland till formationerne. fig. 34. cretaceous chalk at the front of sheet hk10 at hvideklint, overlain by a succession of quaternary deposits including the kraneled, klintholm till, kobbelgård, mid danish till and lolland till formations. midtdanske till fm hk 12 hk 11 kobbelgård fm lolland till fm fig. 35 de stejltstillede skrivekridtskiver hk11 og hk12 i hvideklint, overlejret diskordant af den næsten horisontalt afsatte lolland till formation med en linse af skrivekridt­glacitektonit. per­ son i midten af billedet som skala. fig. 35. the steeply dipping chalk sheets hk11 and hk12 at hvideklint, truncated by the almost horizontally deposited lolland till formation including a lens of chalk-glacitectonite. person for scale in centre of photo. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 36 of 53 www.geusbul let in.org synklinal under overskydningssålen af hk7. hk8 er en relativt kort skive, som blev båret på ryggen af hk9. en endnu mindre og tyndere skive af skrivekridt­glacitek­ tonit er ved sålen af den midtdanske till formation slæbt diskordant hen over næsen af hk8 og hk9. spidsen af skrivekridtskiven hk9 er så kraftigt sheardeformeret, at den er en mellemting mellem en skrivekridt­glacitektonit og en breccieret skrivekridtmoræne (kridt­till, fig. 31, 32). mellem skrivekridtskiverne hk9 og hk10 optræder to skiver med foldede og forskudte enheder af kraneled, klintholm till og kobbelgård formationerne. den øver­ ste af disse skiver er foldet i en overkippet antiklinal, og langs overskydningszonen mellem de to skiver er lagene intensivt deformeret af reverse forkastninger (fig. 33). på ryggen af hk10 optræder mørkt ler oven på skrivekridtoverfladen, som kan henføres til kraneled og klintholm till formationerne. stærkt reverst forskudte, sandede lag består af vekslende indslag af kobbelgård formationen. ved den frontale del af hk10 findes tre bundmoræner, hvor klintholm till formationen er adskilt fra den midtdanske till formation og lolland till formationen ved sandede og lerede indslag af kobbel­ gård formationen (fig. 34). skrivekridtskiverne hk11 og hk12 er to relativt tynde og stejltstående skiver (fig. 35). på grund af shear­ deformationen optræder der både lag og linser af sort ler fig. 36 tynde lag af skrivekridt­glacitek­ tonit i hvideklint, skubbet op over de allerede opskudte skiver hk1–hk12. de tynde lag og slirer af skrivekridt er valset ind i mørkt ler tilhørende kobbelgård formationen ved sheardeformation. fig. 36. thin layers of chalk-glacitectonite at hvideklint, which were squeezed up above the existing pile of chalk sheets hk1–hk12. the thin chalk layers and lenses have been sheared into the dark clay of the kobbelgård formation. fig. 37 slirer og tynde lag af ler i det overskydnings­breccierede skrivekridt i skive hk8 i hvideklint. den oprinde­ lige lagdeling i skrivekridtet er ødelagt, og bjergarten er omdannet til en glacitektonit. fig. 37. lenses and thin layers of clay in the chalk of sheet hk8 at hvideklint, deformed by thrust faulting. the chalk has completely lost its primary sedimentary layering and has been transformed into a glacitectonite. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 37 of 53 www.geusbul let in.org i skrivekridt og lag og slirer af skrivekridt i mørkt ler (fig. 36, 37). fra hk11 mod sydvest tiltager forekomsten af leraflejringer, fortrinsvis fra kraneled og kobbelgård for­ mationerne, hvilket har medført forøget nedskridning og tilmudring af kystprofilet (fig. 38). endnu en skrivekridt­ flage, hk13, er identificeret længere mod sydvest i ret­ ning mod hjelm nakke. her er de tektoniske forhold dog så komplekse, at det ikke giver mening at optegne dem i et tværprofil. et eksempel på denne kompleksitet i defor­ mationsmønsteret ses i forekomsten af mudderdiapirer, som dannes, hvor vandmættede leraflejringer udsættes for glacialtektonisk deformation (fig. 39). glacialdynamisk udvikling pedersen (2012a) har foreslået en glaciodynamisk sekvensstratigrafi for glaciale aflejringer i et terre­ strisk område, som også omfatter aflejringernes glaciale deformation både ved dannelsen af de glacialtek to niske komplekser og selve de breccierede litologier (de såkaldte glacitektoniter, pedersen 1988, 2014; pedersen et al. 2018). til en glacialdynamisk hændelse svarer en glacial­ dynamisk sekvens. basis af en glacialdynamisk sekvens er det dybeste niveau, som påvirkes af den glacial­ dynamiske hændelse. i den komplette glacialdynamiske sekvens er det décollementniveauet for den glacial­ tektoniske deformation, dvs. det dybeste niveau, langs fig. 38 breccierede klaster af skrivekridt i leret langs sålen af overskydnings­ skiverne i hvideklint. på grund af det meget store lerindhold i kraneled og kobbelgård formationerne er klintens sydvestlige del ofte delvist dækkket af mudder. fig. 38. brecciated clasts of chalk in mud along the sole of the thrust faults at hvideklint. the large mud content in the kraneled and kobbelgård formations commonly blurs exposures of the south-western part of the cliff. fig. 39 mudderdiapirisme i hvideklint dannet under glacialtektonisk defor­ mation af vandmættede leraflejringer. blotningen er ca. 2 m høj. foto: tove stockmarr (1996). fig. 39. mud diapirism at hvideklint formed during glacial tectonic deformation of water-saturated clay deposits. the exposure is almost 2 m high. photo: tove stockmarr (1996). http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 38 of 53 www.geusbul let in.org hvilket der sker en glacialtektonisk forskydning. herover følger de enheder, som indgår i det glacial tektoniske kompleks. de kan bestå af prækvartære formationer som f.eks. glacialt deformeret skrivekridt og præ­ eller proglaciale formationer såsom smeltevandsserier. det glacialtektoniske kompleks afskæres af en vinkel­ diskordans, hvorover en glacitektonit dannes ved sålen af en bundmoræne. toppen af den glacialdyna­ miske sekvens er overfladen af den bundmoræne, som blev aflejret under det isfremstød, der forårsagede den glacialdynamiske hændelse (fig. 40). på møn kan man skelne mellem fire glacialdynamiske sekvenser: klintholm, hvideklint, møns klint og aborrebjerg. klintholm glacialdynamiske sekvens kystprofilerne øst og vest for klintholm havn inde­ holder nogle dislocerede enheder af smeltevandsler og ­sand samt nogle diamikte aflejringer, der enten kan være bundmoræneler eller dropstens moræneler (se diskussionen af de diamikte aflejringers dannelse i houmark­nielsen 1994). ved basis af de glaciale aflej­ ringer forekommer skrivekridt og varierende mængder af skrivekridtskiver, der er indarbejdet i de glaciale sedi­ menter. skiverne er skubbet op fra øst mod vest og over­ lejres diskordant af mørkegråt bundmoræneler. denne sekvens betegnes klintholm glacialdynamiske sekvens. den glacialdynamiske hændelse skete under det tidlig­ ste (ristinge) isfremstød i den tidlige til mellemste del af weichsel for ca. 55 000 til 50 000 år siden. klimatisk efterfølges hændelsen af sandnæs mildningen. hvideklint glacialdynamiske sekvens det glacialtektoniske kompleks på hvideklint består af ca. 15 skiver af skrivekridt, der bærer en varie­ rende mængde af glaciale aflejringer på ryggen over top­skrivekridt­inkonformiteten. alle de opskudte skiver skæres af den overliggende midtdanske till for­ mation, hyppigt med en 0,5–1 m tyk, udtværet enhed af skrivekridt­glacitektonit ved bundmorænens sål. denne sekvens betegnes hvideklint glacialdynamiske sekvens, hvis årsag er det dominerende isfremstød, der med et kildeområde i det centrale sverige spredte sig ud til hoved opholdslinjen i danmark, nordtyskland og polen m.v. (nø­isfremstødet). på møn har isfrem­ stødet retning fra nordøst mod sydvest. foruden hvide­ klints glacialtektoniske kompleks kan den anden fase af præstebjerg bakkekomplekset inddrages i denne glacio­ dynamiske sekvens. hændelsen skete i sidste halvdel af weichsel­istiden på det tidspunkt, hvor det seneste gla­ ciale maksimum (lgm) indtrådte. møns klint glacialdynamiske sekvens det glacialtektoniske kompleks på høje møn er arkitek­ tonisk opbygget som en samlet enhed, der fra syd er skubbet op mod nord. enhedens bund befinder sig ved en décollementflade ca. 100 m u.h. (pedersen 2000, 2003; pedersen & gravesen 2009). ved konstruktionen af dybden til décollementfladen er principperne for ba lancerede glacialtektoniske tværprofiler benyttet som beskrevet af pedersen (2005, 2006). over décollement­ fladen kan skrivekridtet ikke betragtes som faststående. selve det glacialtektoniske kompleks har en klassisk opbygning i en proksimal, central og distal del fra syd mod nord. toppen af skrivekridtskiverne og de mellem­ lejrede glaciale aflejringer er afskåret af lolland till for­ mationens bundmoræne, der hyppigt har en medslæbt sål af skrivekridt­glacitektonit. denne sekvens betegnes her som møns klint glacialdynamiske sekvens. den tilhørende glacialdynamiske hændelse tilskrives den hurtigt fremadskridende ungbaltiske is for ca. 18  000 år siden. denne hændelse er tillige ansvarlig for fase 3 og måske 4 i præstebjerg bakkekomplekset (se kapitlet landskabets udformning og dannelse nedenfor). det er den glacialdynamiske dannelse af møns klint, der er ansvarlig for den overordnede udformning af hele møn og specielt kysten ud mod hjelm bugt (fig. 41) samt den overprægede deformation af hvideklint. bundmoræne glacitektonit overskydninger sand og ler décollement skrivekridt d c b a ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ fig. 40 konceptuel illustration af det glacialdynamiske sekvens­ stratigrafiske system (fra pedersen 2012a). a: beliggenheden af den basale décollementzone i toppen af skrivekridtet. b: dannelsen af imbrikerede overskydningsskiver bestående af skrivekridt og glaciale aflejringer. c: dannelsen af glacitektonit ved sålen af den overskridende is. d: afsluttende aflejring af bundmoræne oven på glacitektoniten. fig. 40. the concept of glaciodynamic sequence stratigraphy (after pedersen 2012a). a: location of the basal décollement zone in the upper part of the chalk. b: formation of thrust fault imbricates consisting of chalk and glacial deposits. c: truncation of the glacier sole during formation of glacitectonite. d: final deposition of lodgement till above the glacitectonite. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 39 of 53 www.geusbul let in.org aborrebjerg glacialdynamiske sekvens den nordlige halvdel af det glacialtektoniske kom­ pleks på møns klint er overpræget af en glacialtek­ tonisk opskydning fra nordøst (fig. 42; pedersen 2000). det bedste eksempel på dette er overskydningen af skrivekridtskiven ved vitmunds nakke, som tydeligt er skubbet op fra nordøst mod sydvest, og som også fortsætter lolland till formationen og den midtdan­ ske till formation samt ældre glaciale enheder. ved en analyse af dobbeltfoldningens struktur ses det, at netop aborrebjerg danner kulminationen af en fold på en fold (fig. 43). sekvensen er derfor benævnt aborrebjerg gla­ cialdynamiske sekvens. opskydningen svarer til et recessivt isfremstød fra området nordøst for møn. herfra må den ungbaltiske is under sin tilbagesmeltning have rykket en ’pied­ mont’­formet istunge frem. en sydvestrettet isfront herfra dislocerede de allerede forskubbede kridtskiver. ifølge de gængse modeller for tilbagesmeltningen af den ungbaltiske is (houmark­nielsen 2010, houmark­ nielsen et al. 2017) må dette genfremstød være foregået for ca. 16 000 år siden. landskabets udformning og dannelse geomorfologisk beskrivelse den følgende geomorfologiske beskrivelse er baseret på det geologiske kort og en højdemodel over møn (fig. 44). glacialtektonisk parallelbakket landskab det mest karakteristiske ved høje møn er de langstrakte, parallelle bakkerygge adskilt af langstrakte snævre dalstrøg. bakkeryggene består af de glacialtektonisk opskudte skrivekridtskiver, og talrige steder i store klinteskov træder skrivekridtet frem i overfladen, hvilket især er tydeligt efter kraftigt stormvejr, hvor rodvæltere giver et indsyn i skovbundens bestanddele. i den sydlige del af klinteskovens bakkekompleks er ryggenes retning ø–v. karakteristiske eksempler herpå er gråryg med den langstrakte nælderende nord for gråryg. længere mod nord begynder strygningen at dreje mod øsø–vnv, som det er tilfældet med det dalstrøg, hvori siesøerne ligger, og den nordfor beliggende bakkeryg siesø bjerg. den ås, der går ind i skoven fra dronningestolen, kan ligeledes betragtes tilhørende det omtrent ø–v­strygende bakke­ system. længere mod nord ændrer bakkeryggene mar­ kant retning til sø–nv. enkelte af bakkeryggene længst ude mod klinten drejer helt op til retningen ssø–nnv, og den samme retning kan følges i sølavningerne fra hunesø til liselund. ud fra den glacialtektoniske tolkning af møns klint komplekset kan bakkerne klassificeres som sammen­ satte rygge. ændringerne i deres retning er tolket som resultatet af et istryk først fra sydsydøst, som tilskrives den ungbaltiske is. dette tidligste kompleks af bakke­ rygge blev mod slutningen af istiden overpræget af et isfremstød fra skåne, som med retning fra nordøst har a b fig. 41 blokdiagrammer af møns landskabsdannelse foran den ung­ baltiske is. a: den ungbaltiske is i loben i hjelm bugt. b: blotning af det buede bakkeforløb på møn under afsmeltningen af den ungbaltiske is. bemærk, at bakkerne skæres af smelte vandsstrømme fra gletsjerran­ den langs sydkysten af møn. fig. 41. block diagrams showing the formation of møn’s hilly landscape in front of the young baltic ice. a: the young baltic ice occupying the depression in hjelm bugt. b: exposure of an arcuate, hilly landscape during meltback of the ice. note that the hills are cut by valleys formed by meltwater spillways from the adjacent ice margin along the south coast of møn. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 40 of 53 www.geusbul let in.org a b f1 i denne position opstår aborrebjerg f1 hældning skaber vitmundsnakke f1 f1 f2 fig. 43 dannelsen af aborrebjerg og vitmunds nakke. aborre bjerg er det højeste punkt på møn. at netop dette punkt er så højt, tolkes som resultatet af to forskellige hændelser: a: opskydning af skrivekridtskiver fra syd under det ungbaltiske isfremstød (f1). b: forkast­ ning forårsaget af genfremstød af is akkumuleret i det sydlige sverige (f2). fig. 43. formation of aborrebjerg and vitmunds nakke. aborre bjerg is the highest point of møn and was formed during two separate events: a: superimposed thrust fault deformation by the young baltic ice advance from the south (f1). b: superimposed faulting by the latest ice advance from southern sweden (f2). taglagt vifte ramper antiform stak slotsgavle forlands overskydning maglevandsfald rampe dronningestolen antiform stak gråryg overskydning overskydning fra ønøjydelejet overpræget overskydning ungbaltisk isfremstød fra ssø fig. 42 møns klints strukturelle udvikling (efter pedersen 2000). blå: skrivekridt. brun og orange: glaciale aflejringer. røde linjer: overskydninger. blå pile: istryksretninger fra henholds­ vis ssø og ønø. fig. 42. the structural evolution of møns klint (after pedersen 2000). blue: chalk. brown and orange: glacial deposits. red: thrust faults. blue arrows show the directions of ice push from the sse and ene, respectively. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 41 of 53 www.geusbul let in.org overpræget og til en vis grad reorienteret de allerede anlagte rygge (pedersen 2000). en ø–v­gående orien­ tering af bakkerygge genfindes i det glacialmorfologiske landskab omkring stege nor. selve noret kan betragtes som et morfologisk element, der følger bakkeryggens ret­ ning. de steder, hvor landskabets højeste toppe rager op, tolkes som resultater af, at skiver af skrivekridtunderlaget her allerede var skubbet op fra nordøst. moræneplateauer omkring 70 % af jordarterne på møns overflade består af moræneler, som altovervejende kan henføres til lol­ land till formationen. bortset fra store klinteskov og enkelte andre højere bakkepartier kan denne landskabs­ type med et morænedække karakteriseres som et moræneplateau dannet af en bundmoræne, der giver landskabet en flad, udjævnet terrænform med spredte afløbsløse lavninger. de skyldes dødishuller efter bort­ smeltede isskiver og isklumper, der har været indarbej­ det og begravet i moræneleret under det ungbaltiske eller tidligere isfremstød (fig. 45). under dækmoræneleret ligger en enhed af smelte­ vandssand, ny borre formationen. i landskabet træder dette sand frem i dalstrøg, som det tydeligt ses i dalene øst og vest for stege nor. fra stege nor og sydvestover er sandet udbredt i lavningen omkring damsholte og røddinge sø og i dalstrøget ud mod fanefjord. hatformede bakker og overprægede åse på den vestlige del af møn forekommer en del bakker, som kan tolkes som hatformede bakker eller åse, der har været overskredet af senere isfremstød end dem, der oprindeligt var årsag til pro­ og inglaciale aflej ringer. de proglaciale smeltevandsaflejringers dannelses­ dynamik og deres efterfølgende deformation langs randzonen af isen gennemgås i følgende afsnit. de inglaciale aflejringer er dannet i tunneldale under isen eller i store isdæmmede søer med vandspejl på højde med isoverfladen, men med kanter og bund gennem isen relateret til sprækkezoner i ismassen. efter isens tilbagesmeltning står disse aflejringer frem i landskabet som sand­ og grusbakker, der i den ældre geologiske litteratur omtales som hatformede bakker (berthelsen 1980). det er uvist, hvor mange af vestmøns bakker der kan betragtes som hatformede bakker, men en af dem, præstebjerg, har været genstand for grusgravning, og a b fig. 44 geomorfologiske højdemod­ eller af møn. a: møn set mod nord. b: møn set mod vest. gul­grøn: 0–25 m o.h. brun: 25–50 m o.h. hvid: højder over 50 m o.h. fig. 44. geomorphological elevation model of møn. a: viewed towards north. b: viewed towards west. yellow-green: 0–25 m a.s.l. brown: 25–50 m a.s.l. white: above 50 m a.s.l. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 42 of 53 www.geusbul let in.org under udvindingen af sand og grus kunne bakkens indre struktur studeres (fig. 46; berthelsen 1980). kriteriet for en hatformet bakke er, at den står frem som en isoleret hat i landskabet. det er også en for­ dring, at lagene inde i bakken skal være deformerede og stejlt skråtstillede (denne fordring tilskrives berthelsen (1980), som her har set bort fra talrige ikke dislocerede hatformede bakker f.eks. på langeland). berthelsens (1980) undersøgelse af præstebjerg demonstrerer, hvor komplekse aflejrings­ og deformationsprocesserne i en hatformet bakke kan være. i den første fase blev bakkens materiale aflejret som en proglacial smelte­ vandsslette, der efterfølgende blev foldet op af et isfremstød fra sydøst, antagelig klintholm isfremstødet (fig. 47). om dette isfremstød svarer til ristinge klint till formationens eller klintholm till formationens is kan dog ikke endeligt afgøres uden en detaljeret osl­ eller infrared stimulated luminescence (irsl)­aldersbestem­ melse, men det er tydeligt, at lagserien er deformeret mod nordvest af en baltisk is. den anden fase (fase 2 i fig. 47) er overskridelsen af nordøstisen. umiddelbart inden den midtdanske till formations bundmoræne blev aflejret, blev sandbakken deformeret af et istryk fra nordøst mod sydvest. den tredje fase blev indledt med, at bakkens nordlige side blev stærkt eroderet af en smeltevandsstrøm, som efterfølgende aflejrede en tyk serie smeltevandssand og ­grus. aflejringen skete foran et isfremstød fra øst, som efterfølgende overskred bak­ ken og afsatte en tynd bundmoræne. i den afsluttende fjerde fase blev bakken overskredet af et nyt isfremstød fra sydøst, som foldede bakkens indre lag omkring en nø–sv­strygende foldeakse og overkippede folderne mod nv (fig. 47). i den her fremlagte tolkning foreslås, at de to sidste isfremstød var oscillerende isbevægelser, der kan relateres til den isstrøm, som aflejrede lolland till formationen. endemorænelandskab med smeltevandsslette fra ny borre og sydpå til busemarke spreder sig en større akkumulation af smeltevandssand og ­grus. i dette område har to af møns største grusgrave også ligget, nemlig ny borre grusgrav og en grusgrav i den vestlige udkant af busemarke. området begrænses mod syd af randmorænebakkerne nord og syd for råby sø og fig. 45 grænsen mellem det højtliggende landskab høje møn og det lavere småkuperede bundmorænelandskab set mod sydvest. i horisonten på den fjerne side af hjelm bugt anes de hvide klinter i hvideklint. fig. 45. the boundary between the elevated landscape of høje møn and the landscape towards the west of the lower, undulating till plain. the white cliffs of hvideklint can just be seen in the horizon on the far side of hjelm bugt. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 43 of 53 www.geusbul let in.org a b fig. 46 den hatformede bakke præstebjerg på vestmøn. a: bakken set fra syd. opgravet materiale ses på toppen. b: stejltstillede lag af smelte­ vandssand og ­grus inde i bakken. disse lag skæres øverst og overlejres af moræneler tilhørende lolland till formationen. fig. 46. the hat-formed hill of præstebjerg on western møn. a: the hill viewed from the south, with excavated material at its top. b: vertical layers of meltwater sand and gravel inside the hill, cut and overlain by clayey till of the lolland till formation. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 44 of 53 www.geusbul let in.org busemarke mose med den mest markante bakketop ved kobbelgård. disse bakker udgør et endemoræne­ landskab sammen med de øvrige bakkestrøg, der stryger parallelt med kysten ud mod hjelm bugt. foran rand­ morænebakkerne spredte sig en smeltevandsslette fra busemarke til ny borre, hvorfra afstømningen af smelte­ vand videre nordpå fremeroderede den brede borre sømose lavning. det amerikanske udtryk for et landskab skabt af smeltevandsafstrømning er ’spillway’. dalstrøget op mod maglemose kan betragtes som en lignende dan­ nelse. også lavningen fra nordkysten ind mod borre sømose og maglemose lavningen tjente som ’spillways’ for smelte vandet, da isranden under tilbagesmeltnin­ gen havde en kortvarig stilstand ved kysten af hjelm bugt. smeltevandet løb da især gennem dalstrøget mod maglemose, hvor sand og grus blev omlejret under den senglaciale afsmeltning. en lignende dannelseshistorie har også betinget udviklingen af områderne ved stege nors udmunding mod stege bugt, klostervig, røddinge sø og fanefjord. i alt udgør arealer med smeltevands­ sand knap 10 % af jordarterne i møns overflade. da den ungbaltiske is endelig smeltede tilbage fra møn, fremstod øen som et randmorænelandskab. hjelm bugt udgør et lobeformet bækken, der kan betragtes som den lavning, hvorfra materialet til opbygningen af møn stammer. det marine forland med strandvolde og krumodder for ca. 8000 år siden blev de indre danske farvande oversvømmet af den atlantiske transgression. denne havstigning beløb sig til mere end 20 m, og dens marine aflejringer findes nu i lavtliggende kystområder. de udgør ca. 10 % af jordarterne på møn. borre sømose og maglemose blev på dette tidspunkt omdannet til fjorde, og klostervig udgjorde et smalt stræde mellem nogle kraneled og klintholm till fm midtdanske till fm ældre smeltevandssand ny borre fm lolland till fm 4 3 2 1 fig. 47 dannelsen af præstebjerg. 1: opfoldning af kraneled og klintholm till formationerne under klintholm isfremstødet fra øst­sydøst. 2: genfold­ ning og aflejring af den midtdanske till formation under nø­isfremstødet fra nordøst. 3: smeltevandserosion fra en østlig retning og aflejring af ny borre formationen. 4: genfoldning under det ungbaltiske isfremstød fra sydøst og aflejring af lolland till formationen. delvis efter berthelsen (1979). fig. 47. model for the formation of præstebjerg. 1: folding of the kraneled and klintholm till formations during the klintholm ice advance from east-southeast. 2: superimposed folding and deposition of the mid danish till formation by the ne ice advance from northeast. 3: glaciofluvial erosion by meltwater streams moving from east to west and deposition of the ny borre formation. 4: superimposed deformation due to the young baltic ice advance from south-east and deposition of the lolland till formation. partly from berthelsen (1979). http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 45 of 53 www.geusbul let in.org småøer og holme. på grund af materialestrømmen langs kysten blev disse fjorde og stræder lukket, efterhånden som strandvolde spærrede for havvandets indtrængen. bag strandvoldene dannedes strandsøer, som efter­ følgende groede til og blev til moser med tykke lag af tørv. mest markant blev udviklingen af det marine forland, hvor havstrømmenes materialetransport udbyggede kystområderne. på denne måde blev ulvshale skabt. kortlægningen af strandvolde, der danner krumodder, viser tydeligt en transportretning langs nordkysten af møn fra øst mod vest. ved møns nordspids blev det marine forland udbygget mod nordvest, hvor strand­ voldene danner markante rygge, der hyppigt over­ lej res af en mindre klit. også et mindre område ud for hårbølle havn er dannet som et marint forland. her viser mønsteret af strandvolde, at der blev dannet et vinkelforland, idet transporten skete fra to modsatte retninger. hårbølle forlandet er således både udbyg­ get ved sydgående transport gennem grønsund og ved nordvestgående transport fra hjelm bugt gennem grønsund. transporten af materiale frem til møns nord­ vestlige og sydvestlige hjørner viser tydeligt, at kildeom­ rådet skal findes på møns østside. her er det naturligvis de høje klinter langs møns klint, der har bidraget med erosionsmateriale, specielt fra de mange fjeld­ og jordskred, som beskrives nedenfor. fjeldskred og jordskred fjeldskred forekommer kun få steder i danmark. møns klint er det mest markante område på grund af de høje, stejle skrivekridtklinter (pedersen et al. 1989; nadim et al. 2008; pedersen & gravesen 2009; pedersen 2012b). en anden skredtype er jordskred, som også forekommer hyppigt på møns klint og mange andre steder i landet (nadim et al. 2008). forskellen på de to skredtyper er litologi og terrænstejlhed. fjeldskred opstår, hvor en blok af en helt eller delvis konsolideret bjergart skrider ned fra en stejl klippeside ved et næsten lodret fald, på engelsk betegnet ’rockfall’. jordskred opstår, hvor de geologiske lag består af vandmættede leraflejringer, og hvor skråningens gradient kun behøver at være nogle få grader. på grund af et højt porevandstryk bliver jord­ lagene båret af en mudderfilm og bevæger sig med langsom til moderat hastighed ned ad det skrånende terræn. denne skredtype, som dækker et spektrum fra mudderstrømme til terrasseformede blokudskrid­ ninger, betegnes på engelsk under ét som ’landslides’. fig. 48 fjeldskred ved store taler i januar 2007. flyfoto. fig. 48. landslide at store taler in january 2007. aerial photograph. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 46 of 53 www.geusbul let in.org fjeldskred der er inden for de seneste 100 år kun rapporteret et enkelt tilfælde, hvor et dansk fjeldskred har medført et dødsoffer. det indtraf den 29. juli 1994 fra magle­ vandsnakke lige syd for dronningestolen (pedersen 1994). skreddet var usædvanligt ved at foregå midt på en tør sommerdag. den blok, som faldt ned, var kun på nogle få tusind kubikmeter (2500–3000 m3, peder­ sen 1994), men da faldet skete på et tidspunkt, hvor talrige turister færdedes på stranden, var risikoen for et fatalt uheld stor. en person blev dræbt under skred­ det, og siden har klinten på det pågældende sted været genstand for omhyggelig overvågning. imidlertid er det vanskeligt at forudse, hvor et kommende fjeldskred vil ske. en registrering af de seneste 50 års fjeldskred har vist tegn på, at et større skred forekommer omkring hvert femte år (tabel 1). det seneste større skred, hvor 100 000 m3 materiale spredte sig 300 m ud fra kysten, indtraf i januar 2007 ved fjeldpartiet taleren (fig. 48; pedersen 2007). samme år skete der flere andre skred langs møns klint og andre steder i landet. årsagen til disse skred var et ualmindelig regnfuldt efterår, hvor den gennemsnitlige grundvandsstand stod ca. 1 m over normal vandstanden. derfor var porevandstryk­ ket meget højt i de øverste jordlag, hvilket betød, at brudstyrken var sænket. med nedsat brudstyrke er hundevangsfaldet fruerstuefaldet skredrenden tragten sandfaldet maglevandsfaldet maglevandsfaldet sandskredsfaldet sækkepibefaldet gukkenhule dronningestolen 24 7 33 11 33 15 5 4 21 29 29 33 32 39 46 35 24 48 6 34 41 33 24 36 36 29 29 29 33 38 34 34 40 41 35 35 36 32 44 21 21 17 kyst ca. 1999 kyst ca. 1889 200 m n s erosion i m35 fig. 50 erosion og tilbagetrækning af møns klint over tid på grund af skred og erosion, vist ved forskellen mellem kystens beliggenhed i 1899 og ca. 1999. fig. 50. erosion and retreat of møns klint over time by landslides and erosion, shown as the calculated difference between the coastlines in 1899 and c. 1999. fig. 49 efter fjeldskred i klinten er fossil­ jagt meget populært i det nedskredne skrivekridt. foto: peter moors. fig. 49. after rockfalls from the cliffs, fossil hunting in the displaced chalk material is very popular. photo: peter moors. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 47 of 53 www.geusbul let in.org tilbøjeligheden til sprækkedannelse større, og når der først er dannet sprækker, er grundlaget for en udskrid­ ning skabt. hver gang et større fjeldskred er sket, er der dan­ net en halvø af skrivekridt (fig. 49). de friske klumper af skrivekridt er et yndet mål for fossiljægere. normalt afspærres området lige neden for fjeldskredfladen, da efterskred på et par kubikmeters størrelse hyppigt sker op til et par døgn efter et stort kollaps. en opgørelse af erosionsraten kan give et overordnet udtryk for skredforholdene. ved at sammenligne kystlin­ jen langs klinten fra 1899 med kystlinjen opmålt 100 år senere får man et godt billede af erosionsraten forskellige steder langs kysten (fig. 50). den gennemsnitlige erosions­ rate er knap 30 cm om året, hvilket er omkring dobbelt så meget som den almindelige erosion langs andre danske klintkyster. de steder, hvor erosionsraten er højest, 44–49 cm/år, er dog ikke ud for de stejle skrivekridt profiler, men derimod ud for de store fald mellem stejlklinterne. således er sandfaldet mellem sommerspirklinten og græderen syd for dronningestolen et udsat sted, idet der forekom­ mer et kildevæld, og sandmasserne er relativt løse med stor tendens til vedvarende sandskred. jordskred jordskred, hvor lerrige klinter skrider ud, forekommer almindeligt langs de fleste danske klintkyster. på møn er det især den nordlige del af møns klint, der er udsat for denne skredtype. de største kendte skred skete i forrige århundrede, først i 1905 efter en større storm­ flod i østersøområdet, hvor omkring 4 ha af liselunds parkområde skred ud i østersøen. femten år senere blev jordskreddene ved liselund igen aktiveret, dog kun med skred på den halve størrelse, og i 1927 var der igen aktive udskridninger samme sted. endnu i dag kan man se terrassemorfologi og slugter, som skyldes de tidligere jordskred (hintze 1904, 1920). ved foden af liselunds kystskrænt ser man årsagen til de mange jordskred. her træder det blågrå issøler fra kobbelgård forma­ tionen frem i strandplanet. det er på dette blågrå ler, at fig. 51 jordskred på stranden neden for liselund slot med ler fra kobbelgård formationen og nedstyrtede træer. fig. 51. landslide on the beach below liselund castle with clay from the kobbelgård formation and fallen trees. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 48 of 53 www.geusbul let in.org hele udskridningen foregår. skreddene er stadig aktive, hvilket ses af de mange væltede træer, som ligger på stranden, der er domineret af flintrigt grus og sten (fig. 51). ikke kun ved liselund forårsager leret fra kobbel­ gård formationen jordskred. i efteråret 2007 skete der en cylindrisk udskridning ved abildgaards fald, hvorved hele strandplanet ved foden af skreddet blev løftet 6 m op over havoverfladen (fig. 52). hele kystklinten længere nordpå fra liselund via hellehavn nakke til pomlerende er domineret af jordskred, der forskyder de kvartære lagserier, som karakteriserer møns klints nordlige del. english summary the geological map sheet møn covers the island of møn, the smaller neighbouring islands langø, lindholm and nyord, and adjacent parts of sjælland and lolland. it comprises the geodetic map sheets 1511 i and 1511 ii and areas on bordering sheets. møn is surrounded by the baltic sea with the bay of hjelm bugt to the south, the straits of grønsund and ulvsund to the west, and the bays of stege bugt and fakse bugt to the north. møn is divided into three geomorphological areas, namely the high, hilly landscape of høje møn to the east, the hummocky to parallel ridge landscape to the west and the areas of marine deposits around nyord and ulvshale. the composite ridge landscape of høje møn constitutes a glacial tectonic complex comprising four individual glacial dynamic sequences, with the hill aborre bjerg as the highest point (143 m a.s.l.). the parallel ridge hills consist of thrust­fault­displaced chalk sheets with superimposed glacial deposits. the thrust sheets are up to 80 m thick, of which 60 m constitute maa strichtian chalk. the vertical displacement of the thrust sheets is about 150 m measured from the pri­ mary, undeformed pre­quaternary surface located 25–30 m below sea level. the pre­quaternary surface table 1 større fjeldskred og jordskred ved møns klint mellem 1801 og 2015§ årstal dato lokalitet størrelse klasse† 1801 ved taleren 1868 24/12 dele af dronningestolen 3 000 000 m3 3 1899 sækkepibe fald 100 000 m3 3 1904 31/12 nord for sommerspiret og ved gukkende hule 1905 05/11 liselund lerskred 4 ha 3 1910 13/03 hundefangsklinten 1914 nytårsnat ved dronningestolen og forchhammers pynt 1920 14/09 liselund, lerskred 4 tdr. land 3 1927 05/11 større skred ved liselund 1927 22/11 slotsgavlene, pomlerende, nonnebænken, stubberup have 1928 11/05 hundefangsfaldet 20–30 tdr land 3 1929 sydlige del af dronningestolen 1939 25/01 puggaards pynt 2–300 m halvø 3 1940–46 gukkende hule, del af vidmunds nakke, forchhammers pynt 1948 15/01 dronningestolen 1952 24/01 røde udfald med dele af vitmunds nakke og puggaards klint halvø 500 m 3 1958 vitmunds nakke 1963 maglevandsfaldet 1970 store stejlebjerg 2 1979 gavl af sommerspiret 2 1980 hylledalsfald 1 1988 13/01 sommerspiret 1 1994 29/07 sydlige del dronningestolen. fransk turist dræbt kl. 12:30 3000 m3 1 1998 04/03 freuchens pynt, 100 m ud i havet 3 2003 05/07 store stejlbjerg, kl. 8:50 12 000 m3 2 2004 16/01 nordlige del af jættebrinken, kl. 16 4000 m3 2 2004 17/03 sommerspirpynt 2000 m3 2 2007 27/01 store taler. nat 100 250 m3, halvø 300 m 3 2007 02/03 freuchens pynt, formiddag 25 000 m3, halvø 150 m 3 2007 30/09 abildgaards fald, nat. lerskred, skredtå, strandbred løftet 6 m op 2 2009 15/10 nylands nakke, nat. tør sensommer slår om til nattefrost. over­ hæng fjeldskred 5000 m3 1 2015 05/12 græderen faldet efter orkanen bodil 1 § tabellen omfatter både fjeldskred af skrivekridt og nedskredet kvartært materiale. landslides along møns klint between 1801 and 2015. comprises both rockfalls of cretaceous chalk and landslides of quaternary deposits. † 1: mindre. 2: middel. 3: stort. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 49 of 53 www.geusbul let in.org consists of chalk of late maastrichtian age, which forms a carbonate platform in the subsurface of møn about 27 m below sea level. chalk displaced by glacial tectonics is not restricted to høje møn but also appears in smaller thrust sheets and rafts in the small­ridged landscape around stege nor. in the chalk sheets along møns klint most of the late maastrichtian succession is exposed. cliff sections with chalk are also exposed at hvideklint along the south coast of the island. however, here the glacial tectonic shear deformation has commonly altered the lithology into a chalk glacitectonite. the oldest quaternary units deposited on the pre­quaternary unconformity are saalian till as well as sand and clay from the eemian interglacial. these units are overlain by early weichselian sand. the next qua­ ternary succession, the ristinge klint till formation, was deposited during the ristinge ice advance in the early middle weichselian about 55 000–50 000 years ago. then followed the kraneled formation (new formation) consisting of fluvial and lacustrine deposits. the follow­ ing klintholm till formation (adjusted formation) was deposited during the klintholm ice advance 35  000– 32 000 years ago. the klintholm till formation is overlain by a more than 10 m thick unit of greyish glaciolacustrine clay with dropstones. glaciofluvial sand with thin­layered intercalations of laminated mud and diamictites of the kobbelgård formation (new formation) are related to this unit and interpreted as deposited in a huge, partly ice dammed lake covering a large part of the present baltic sea and the southern part of kattegat 32 000 to 28 000 years ago. the kobbelgård formation is overlain by sand and gravel of the stubberup have formation (new formation) and tills of the mid danish till formation deposited by the advance of nordøstisen (the north­east ice) from central sweden about 23 000–20 000 years ago. relatively shortly after the ne­ice had melted away, the ungbaltiske is (young baltic ice) advanced from the eastern part of the baltic area. north­directed antaget position af vandkant før jordskred hævet klintfod ved jordskred hævet strandbred vandkant hævet ved jordskred september 2007 fig. 52 jordskred ved abildgaards fald i den nordlige del af møns klint, september 2007. de tre linjer viser jordskreddets front, hævede strandvolde og den nuværende strandbred med antaget forlængelse. fig. 52. landslide at abildgaards fald in northern part of møns klint in september 2007. the toe of the slide displaced the beach plane. the three lines mark the elevated cliff foot, the elevated former shoreline and its recent position with an extrapolation. http://www.geusbulletin.org pedersen & gravesen 2021: geus bulletin 48. 8293. https://doi.org/10.34194/geusb.v48.8293 50 of 53 www.geusbul let in.org compressive deformation during his advance created the glacial­tectonic complex of møns klint including the new unit ‘møns klint glacialdynamiske sekvens’. in the southern part of the complex, a steeply inclined imbri­ cated fan was formed; towards the foreland to the north the thrust faults became gently dipping and the tip­zone of thrusting is located under the landslides at liselund. the composite ridges form a characteristic hilly land­ scape with elongate crests trending e–w. the curved coastline along hjelm bugt was formed by a glacial lobe, north of which a push moraine was built up. a number of spillways striking radially northward from the lobe were formed by meltwater discharged from its glacier ports, including the borre, maglemose and røddinge depressions. deposition of sand and gravel of the ny borre formation (new formation) took place at this time. during the advance of the young bal­ tic ice over southern denmark to the eastern jutland sta­ tionary line, a relatively thin lodgement till of the lolland till formation was deposited, which is rich in chalk due to its truncation of the upthrusted chalk sheets. towards the end of the weichselian glaciation c. 17 000 years ago, the young baltic ice melted back, leav­ ing a residual ice cap in skåne from where a recessive ice advance towards south­west reached møns klint, resulting in superimposed glacial tectonic deformation. during the late weichselian, freshwater lakes in the hjelm, tøvelde and høje møn areas were filled by clay and gyttja, with deposition that continued into the holocene. during the holocene, the former spillways were turned into fjords during the atlantic transgression. marine deposits mirroring the littorina sea are thus found in maglemose and borre sømose. after the atlan­ tic transgression had established a sea level more or less corresponding to that of the today, accretion of marine forelands and formation of a spit system started. in particular, this is the case for the areas of ulvshale and nyord. at the same time vegetation migrated out into the numerous fjords, and peat began to accumu­ late. the last phase of sedimentation is confined to the formation of beach ridges in the coastal areas, typically covered by aeolian dunes, as can be seen on the coast at klintholm havn and råbylille as well as along the north­eastern coast of ulvshale. tak en stor tak skal rettes til tove stockmarr og gerald hyde for tilladelse til at anvende deres specialeopgaver, herunder figurer og fotos. gunver krarup pedersen takkes for mange konstruktive faglige kommentarer til en tidligere version af teksten. de to eksterne bedømmere helena alexanderson og nikolaj krogh larsen takkes for omhyggelig gennemgang af teksten og forslag til forbedringer. ansvaret for det geologiske kort og den endelige tekst er forfatternes. additional information funding funded by geus. additional files the møn geological map sheet is supplied as an additional file at https://doi.org/10.22008/fk2/tmfh5w references aaris­sørensen, k., petersen, k.s. & tauber, h. 1990: 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(red.): danmarks geologi fra kridt til i dag, 31–67. aarhus: geologisk institut, aarhus universitet ukkonen, p., arppe, l., houmark­nielsen, m., kjær, k.h. & karhu, j.a. 2007: mis 3 mammoth remains from sweden – implications for faunal history, palaeoclimate and glaciation chronology. quaternary science reviews 26, 3081–3098. https://doi.org/10.1016/j.quascirev.2007.06.021 ødum, h. 1933: marint interglacial paa sjælland, hven, møn og rügen. danmarks geologiske undersøgelse iv række 2(10), 44 sider. https:// doi.org/10.34194/raekke4.v2.6980 http://www.geusbulletin.org https://doi.org/10.34194/raekke3.v36.6941 https://doi.org/10.34194/raekke3.v36.6941 https://doi.org/10.37570/bgsd-2006-54-01 https://doi.org/10.1016/j.cretres.2013.08.006 https://doi.org/10.1016/j.quascirev.2007.06.021 https://doi.org/10.34194/raekke4.v2.6980 https://doi.org/10.34194/raekke4.v2.6980 kortbladsbeskrivelse, geologisk kort over danmark, 1:50 000, møn dele af 1511 i, 1511 iv og 1512 ii sammendrag introduktion topografi metode kortlægningens historie prækvartære aflejringer skrivekridt fra sen kridt glaciale og interglaciale aflejringer data og tolkning ældre kvartære aflejringer: saale, eem, tidlig weichsel og tidlig mellem weichsel (mis 6, 5e, 5a-5d, saale (mis 6) eem (mis 5e) tidlig weichsel og tidlig mellem weichsel (mis 5a-5d, mis 4) mellem og sen weichsel (mis 3 og 2) ristinge klint till formation (mis 3) kraneled formation (mis 3) klintholm till formation (mis 3) kobbelgård formation (mis 3) stubberup have formation (mis 3) midtdanske till formation (mis 2) ny borre formation (mis 2) lolland till formation (mis 2) sen weichsel (senglacial, mis 2) udviklingen i sen weichsel (senglacial) holocæne aflejringer (postglacial, mis 1) ferskvandsaflejringer marine aflejringer klitdannelser udviklingen gennem holocæn lagenes optræden og strukturelle forhold de geologiske tværsnit baseret på boringer og andre kortlægninger tværsnit a-b tværsnit c-d tværsnit e-f tværsnit g-h skrivekridtprofiler i kystklinter møns klint hvideklint glacialdynamisk udvikling klintholm glacialdynamiske sekvens hvideklint glacialdynamiske sekvens møns klint glacialdynamiske sekvens aborrebjerg glacialdynamiske sekvens landskabets udformning og dannelse geomorfologisk beskrivelse glacialtektonisk parallelbakket landskab moræneplateauer hatformede bakker og overprægede åse endemorænelandskab med smeltevandsslette det marine forland med strandvolde og krumodder fjeldskred og jordskred fjeldskred jordskred english summary tak additional information references figures fig 1: index map showing position of the map sheet møn. see caption for full description. fig 2: top of chalk below the quaternary deposits. see caption for full description. fig 3: photo of pale grey chalk with black flint layers. see caption for full description. fig 4: palaeogeographic map of the late cretaceous sea in northern europe. see caption for full desc fig 5: stratigraphic divisions. see caption for full description. fig 6 a: schematic stratigraphic log and a locality map. see caption for full description. fig 6 b: schematic stratigraphic log and a locality map. see caption for full description. fig 7 a: eemian deposits at hjelm nakke. see caption for full description. fig 7 b: eemian deposits at hjelm nakke. see caption for full description. fig 8: composite section showing the thicknesses and horizontal distribution of the lithostratigraph fig 9: photo of the ristinge klint till formation at store stejlebjerg, møns klint. see caption for fig 10: photo of the ristinge klint till formation. see caption for full description. fig 11: photo of the klintholm till formation. see caption for full description. fig 12: photo of the kobbelgård formation. see caption for full description. fig 13: photo of the kobbelgård formation. see caption for full description. fig 14 a: photo and sedimentological log of the stubberup have formation. see caption for full descr fig 14 b: photo and sedimentological log of the stubberup have formation. see caption for full descr fig 15: photo of the mid danish till formation. see caption for full description. fig 16: sedimentological log of the ny borre formation. see caption for full description. fig 17: photo of the ny borre formation. see caption for full description. fig 18: photo of the lolland till formation. see caption for full description. fig 19: photo of the lolland till formation. see caption for full description. fig 20: photo of late weichselian – holocene deposits. see caption for full description. fig 21: photo of late weichselian freshwater deposits. see caption for full description. fig 22: section through borre sømose. see caption for full description. fig 23: photo of holocene beach ridges. see caption for full description. fig 24: photos of holocene beach ridges. see caption for full description. fig 25: geological profile g–h. see caption for full description. fig 26: photo of exposed elevated chalk sheets. see caption for full description. fig 27: historical illustration of møns klint profile. see caption for full description. fig 28: the structure of hvideklint. see caption for full description. fig 29: photos of folding at hvideklint. see caption for full description. fig 30: photo of thrust fault at hvideklint. see caption for full description. fig 31: photo of thrust fault at hvideklint. see caption for full description. fig 32: photo of deformed chalk at hvideklint. see caption for full description. fig 33: photo of folded and thrust-displaced layers at hvideklint. see caption for full description. fig 34: photo of cretaceous chalk at hvideklint. see caption for full description. fig 35: photo of steeply dipping chalk sheets. see caption for full description. fig 38: photo of brecciated clasts of chalk. see caption for full description. fig 39: photo of mud diapirism at hvideklint. see caption for full description. fig 40: conceptual model of glaciodynamic sequence stratigraphy. see caption for full description. fig 41: block diagrams showing the formation of møn’s hilly landscape. see caption for full descript fig 42: structural evolution of møns klint. see caption for full description. fig 43: formation of aborrebjerg and vitmunds nakke. see caption for full description. fig 44: geomorphological elevation model of møn. see caption for full description. fig 45: photo of the boundary between elevated høje møn and the lower, undulating till plain. see ca fig 46: photos of the form and structure of præstebjerg hill. see caption for full description. fig 47: model for the formation of præstebjerg. see caption for full description. fig 48: photo of the 2007 landslide at store taler. see caption for full description. fig 49: photo of fossil hunters following a rockfall. see caption for full description. fig 50: schematic showing erosion and retreat of møns klint. see caption for full description. fig 51: photo of landslide remnants. see caption for full description. fig 52: photo of landslide at abildgaards fald. see caption for full description. table 1: record of large landslides at møns klint between 1801 and 2015. see caption for full descri monograph rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 1 of 66 the pge-au mineralisation of the skaergaard intrusion: precious metal minerals, petrography and ore genesis nikolay s. rudashevsky1, troels f.d. nielsen2* , vladimir n. rudashevsky1 1cnt instruments llc, st. petersburg, russia. 2department of mapping and mineral resources, geological survey of denmark and greenland (geus), copenhagen, denmark. abstract the skaergaard pge-au mineralisation, alias the platinova reef, is hosted in a series of mineralisation levels within a suite of bowl-shaped macrorhythmic layers in the upper middle zone of the skaergaard intrusion. the intrusion is exposed 68°n in east greenland. the occurrence defines its own type due to its exceptional structure and mineralogy. a wealth of mineralogical data is available in laboratory reports for individual samples and in peer-reviewed publications, but none of these account for the lateral and stratigraphic distribution of pge and au parageneses in the gabbros of the intrusion. in this study, we collate and describe the mineralogical data for the first-formed pge-rich and last-formed gold-rich mineralisation levels and integrate these with petrogenetic models. recovery of >4000 grains of precious metal phases allow a detailed study of their distribution and compositions throughout the mineralisation, re-equilibration during cooling, inter-grain relationships and relationships to cu-fe sulphides and the gabbroic host rocks. the sulphides are dominated by bornite, chalcocite and minor chalcopyrite. all other sulphides, such as pentlandite, are very rare. fifty-four different precious metal phases are identified in this study, and include the new ima approved minerals skaergaardite (pdcu), nielsenite (pd3pb) and naldrettite (pd2sb). precious metal phases include (1) intermetallic compounds and alloys of cu and pd; (2) intermetallic compounds and alloys of au and cu (ag); (3) sulphides of pd, cu (ag, cd, hg, tl); (4) arsenides of pd (pt, ni) and (5) intermetallic compounds of pd, cu with sn, pb, te (sb, bi). skaergaardite (pdcu) is the dominant pge mineral in the lower and main pge mineralisation level (pd5). it is accompanied at the western margin of the intrusions by the sulphides vasilite (pd16s7) and vysotskite (pds) but is rare at the eastern margin, which is dominated by plumbide zvyagintsevite (pd3pb). gold phases include a suite of intermetallic compounds and alloys from aucu3 to native gold and are dominated by tetra-auricupride (aucu). gold is concentrated in the tops of individual mineralisation levels and in the uppermost precious metal–bearing mineralisation level, followed by stratiform cu-rich mineralisation levels. precious metal parageneses demonstrate formation and re-equilibration from liquidus to subsolidus temperatures and control by local geochemical environments. the mineralisation is syn-magmatic and the result of fractionation and evolution in the remaining bulk-silicate liquid and crystal mushes. fractionation led to sulphide saturation and formation of immiscible sulphide melt droplets. this was followed by reaction with mush melts and re-equilibration to lower temperatures, first under the roof and subsequently after slumping to the floor in mushes of macrorhythmic layers. droplets of sulphide melt formed between 1030–1050°c and trapped precious metals. the subsequent reaction between sulphide melt and interstitial fe-rich immiscible melt at c. 1015°c, and redistribution to coexisting melt and fluid, led to the separation of pge, au and cu and their up and inward transport. magmatic fluids as well as volatile-rich residual silicate melts were retained in gabbros at the margins and resulted in precious metal parageneses in equilibrium with hydrous low-temperature silicate parageneses. *correspondence: tfn@geus.dk received: 04 nov 2021 revised: 16 june 2022 accepted: 13 dec 2022 published: 27 july 2023 keywords: skaergaard intrusion, precious metal minerals, syn-magmatic, crystal mushes, south-east greenland abbreviations bse: backscattered electron bq: drill core standard, diameter 36.5 mm. ecd: equivalent circle diameter emp: electron microprobe qfm: quartz-fayalite-magnetite ƒo2: oxygen fugacity hfse: high field strength elements hrxct: high resolution x-ray computed tomography hs: hydro-separation hz: hidden zone ima: international mineralogical association iss: intermediate solid solution la-iсp-ms: laser ablation inductively coupled plasma mass spectrometry lld: liquid line of descent ls: layered series lz: lower zone mbs: marginal border series mz: middle zone pge: platinum group elements sem: scanning electron microscope t: temperature tof: toe of forbindelsesgletscher ubs: upper border series uz: upper zone geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: rune berg-edland larsen (ntnu, norway) reviewed by: federica zaccarini (montanuniversität leoben, austria), eduardo teixeira mansur (geological survey of norway). funding: see page 61 competing interests: see page 61 additional files: none provided https://doi.org/10.34194/geusb.v54.8306 https://orcid.org/0000-0002-4932-3869 https://orcid.org/0000-0001-9521-7162 mailto:tfn@geus.dk rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 2 of 66 www.geusbul let in.org 1 introduction a potentially economic palladium and gold mineralisation was discovered in the skaergaard intrusion in the late 1980s. the 56 ma layered gabbro intrusion is located 68°n on the east coast of greenland and has played a fundamental role in the development of igneous petrology (e.g. wager & brown 1968). more surprising, and of geological importance, is the 1986 discovery of the large precious metal mineralisation. the estimated resource is >30 million ounces of platinum group elements (pges) dominated by palladium (pd) and c. 9 million ounces of gold (au) in an estimated ore reserve of 1.8 billion tons. it is a surprisingly large amount of precious metal in an intrusion with an original total volume of no more than c. 300 km3. pge and au deposits occur in a succession of mineralisation levels hosted in five macrorhythmic layers (nielsen et al. 2019a). mineralisation level pd5 in the lowermost of these layers is approximately 5 m thick at lower and upper cutoff grades of 0.7 ppm precious metals. assays are found in watts, griffis & mcouat ltd (1991). pd5 is estimated to host c. 450 million tons of ore rich in palladium with an average precious metal content of c. 2 ppm over a width of 5 m (nielsen et al. 2005; kuo 2007). the cutoff is the assumed minimum grade required for economic exploitation of the resource. material with <0.7 ppm precious metal is considered waste. the mineralisation was initially referred to as the platinova reef (bird et al. 1991) and named after the exploration company that discovered the mineralisation. it is not a reef in any common sense of the word, but a series of stratiform, bowl-shaped layers of gabbro enriched in precious metals (nielsen et al. 2015). the mineralised gabbros show remarkable and systematic lateral and stratigraphic variations in ratios between pge and au and between precious metals and copper (andersen et al. 1998). due to its structure and the distribution of the precious metals, the mineralisation was defined as its own type and is thus referred to as “skaergaard-type” by miller & andersen (2002). the mineralisation is here, and following nielsen et al. (2015, 2019a), referred to as the skaergaard pge-au mineralisation. this term avoids conceptional associations to other entirely pge and, for example, nickel-rich (ni) and/or copper-rich (cu) or chromite-related reeftype mineralisations in layered mafic intrusions. examples include sill complexes in western siberia (russia), sudbury-type occurrences (canada), the merensky reef and platreef mineralisations in the bushveld igneous complex in south africa and the j-m reef in the stillwater intrusion, montana, usa (e.g. barnes et al. 2017). unlike most of these pge-rich deposits, the precious metal mineralisation in the skaergaard intrusion is very poor in platinum, copper, nickel and sulphur. the sonju lake (minnesota, usa; maes et al. 2007) and the rincon del tigre (bolivia; prendergast 2000) deposits are also referred to as skaergaard-type. the significant differences in styles of mineralisation and marked paragenetic mineralogical variations reflect the diversity in genesis of pge-rich ores. classic petrogenetic models, inspired by the pge occurrences in, for example, the bushveld igneous complex, argued for orthomagmatic origins and formation of immiscible sulphide droplets that equilibrate with bulk liquid (e.g. naldrett 2004, 2011). the sulphide droplets scavenge the precious metals due to very high distribution coefficients for precious metals between sulphide and silicate liquids. the high-density droplets subsequently accumulate at or in the crystallisation front in reefs of layered mafic intrusions. sulphide saturation in bulk liquid can be caused by many different processes (barnes et al. 2017), such as magma mixing, contamination and fractional crystallisation. alternatively, pge-rich ores were proposed to be the result of redeposition of precious metals, scavenged from already crystallised gabbro and transported in late residual melts or fluids in solidifying magma chambers (e.g. boudreau & meurer 1999; boudreau 2019). neither of these models, however, explain the extremely well-controlled stratigraphic distribution of pge and au in the layered gabbros of the skaergaard intrusion (nielsen et al. 2015), or the lateral variations in the precious metal parageneses (nielsen et al. 2019a). a concordant layer that can be followed across a layered gabbro intrusion is in a classic orthomagmatic model understood to be a time marker. such a layer is formed at a specific stage in the evolution of the bulk magma, and the layer would be expected to have a composition that reflects the evolutionary stage of the bulk liquid. in any given concordant layer, primary and preserved sulphide droplets should have a near-constant composition. that is not observed in the skaergaard pge-au mineralisation. nielsen et al. (2005, 2015, 2019a) therefore developed a petrogenetic model for skaergaard-type mineralisations. the proposed model combines previously suggested processes into a more complex model that includes a succession of orthoto late-magmatic processes. of fundamental importance was the recognition of the role played by in situ sulphide saturation in melts of crystal mushes, reaction with fe-rich immiscible mush melts and the coexistence of immiscible silicate melts, immiscible sulphide melts, and hydrous fluids in crystal mushes below the roof and in the intermittent https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 3 of 66 www.geusbul let in.org floor of the layered gabbros. the importance of mushstage processes is best illustrated by the voluminous mantles and rims on liquidus cores of plagioclase in namur et al. (2014; see also nielsen et al. 2019b). in this study we compile and describe available information on stratigraphic and lateral variations in the parageneses and compositions of the precious metal phases in the type-defining skaergaard intrusion. the compilation of mineralogical data is based on numerous laboratory reports and peer-reviewed studies, and can serve as reference for other skaergaard-type mineralisations. the lateral variations in mineralogy (nielsen et al. 2019a) and bulk-rock composition (nielsen et al. 2015) are explicit in lithological layers that are traced across the intrusion. no single profile or combined profile (e.g. holwell et al. 2016) is representative and cannot be used as the basis for modelling of the type of processes responsible for all of the skaergaard pge-au mineralisation. it is also noted that no simple relationship exists between the observed distribution of all precious metal minerals and the evolution of the contemporary bulk liquid (nielsen et al. 2019a). we subsequently integrate the mineralogical information with the proposed petrogenetic model of nielsen et al. (2015). the integration of data and model provides constraints for the succession of processes responsible for the mineralisation, from the initiation of bulk-liquid crystallisation to low-temperature (t) re-equilibration of the precious metal parageneses. the initial accumulation of the precious metals was apparently a short-lived event while the melt cooled from c. 1050°c to 1030°c, whereas the redistribution of precious metals and re-equilibration of parageneses continued long after initial, and strictly concordant accumulation of the precious elements in the intermittent floor of the intrusion. we underpin that the skaergaard-type pge-au mineralisation is a distinct type of syn-magmatic mineralisation and the result of a complex interplay between processes in crystal mushes over a wide temperature range between liquidus and solidus, and a potential source of critical commodities in comparatively small and evolved tholeiitic intrusions of limited size. an introduction to the skaergaard intrusion and its mineralisation is given in chapter 2. methods used in the original mineralogical investigations are briefly described in chapter 3. in chapter 4, we synthesize the precious metal mineralogy, phase compositions, phase relations and parageneses of the skaergaard mineralisation including descriptions of the gabbroic host (4.1) and ore minerals (4.2), followed by the petrographic relations between host rocks, sulphides and precious metal phases (4.3). in chapter 5, we identify and discuss constraints for the development of the integrated mineralogical and petrogenetic mineralisation model, including the composition of bulk liquid at mineralisation (5.1), timing and temperature of sulphide saturation (5.2), evidence for a high-t origin of initial distribution of precious metals and the first-formed precious minerals and phases (5.3). we then consider the composition of coexisting fluid and reactions (5.4), low-t re-equilibration of the precious metal minerals (5.5) and petrographic information relevant to ore beneficiation (5.6). finally, the combined petrogenetic and mineralogical mineralisation model is presented in chapter 6. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 4 of 66 www.geusbul let in.org 2 the skaergaard intrusion the layered gabbros of the skaergaard intrusion (fig. 1) have been studied in significant detail for more than 75 years (brooks 2011). it is due to exposure and preservation of the gabbros in the arctic climate that these rocks are regarded as a foremost natural laboratory for the study of the fractionation of basaltic magma and processes in shallow crustal magma chambers (e.g. wager & deer 1939; wager & brown 1968; mcbirney 1996; irvine et al. 1998; nielsen 2004; tegner et al. 2009; holness et al. 2011; nielsen et al. 2015). the intrusion was emplaced 56.02 million years ago during the opening of the north atlantic (larsen & tegner 2006; wotzlaw et al. 2012). the intrusion has a box-like shape with the surface dimensions of c. 11 x 7.5 km and an original average height of the magma column of c. 3.8 km. the magma chamber had a volume of approximately 300 km3 (nielsen 2004; svennevig & guarnieri 2012) and was filled over a short period of time (annen et al. 2022). uz ubs mbs uttental sund 31°45’ 31°40’ 31°35’ 68° 13’ 68° 10’ mz lz mz uz forbindelsesgletscher mbs mbs 0 1 2 (a) km 90-18 90-10 tof 90-24 90-23a enewsw sh sh 2 km mbsmbs s. l. uzc uzb uza mz lzc lzb lza+hz ubsubs mbsmbs fault 0 500 1000 1500 2000 2500 3000 lza’ aug+ an59 mt+ an52 wo+ an32 wo+ an33 ap+ an39 ol+ an44 olan50 mt+ an52 aug+ an59 an50 an44 ap+ an40 lz a* lz b* lz c* m z* m za * m zb * lza hz lzb’ lzb lzc’ lzc mz’ mz uza’ uza uzb’ uzb uzc’ uzc la ye re d se ri es u pp er b or de r s er ie s greenland skaergaard intrusion (b) (c) 90-22 st ra ti gr ap hi c th ic kn es s (m ) no vertical exaggeration wagertoppen midnat chip line zones and subzones in ls, mbs and ubs drill cores centre of skaergaard pge-au mineralisation mineralisation in (a) inferred boundaries are stippled internal boundary contact cross-section sea ice later ma�c intrusion precambrian basement flood basalts mineralisation in (b) and (c) ls mbs ubs kraemer ø fig. 1 overview of the skaergaard intrusion. a: simplified geological map slightly modified from nielsen et al. (2015, 2019a), showing the layered series (ls), the marginal border series (mbs) and the upper border series (ubs). the subzones of the ls include the lower zone (lz), middle zone (mz), and upper zone (uz). in red, the collar locations for drill cores 90-10, 90-18, 90-23a and 90-24 and sampling site ‘toe of forbindelsesgletscher’ (tof). b: wsw–ene section (red stippled line in (a)) through the intrusion with subdivisions as in (a), and subzones a, b and c in lz and uz. c: correlation between zones and sub-zones of ls, mbs and ubs (modified from salmonsen & tegner 2013). https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 5 of 66 www.geusbul let in.org the western and northern margins of the intrusion are chilled to precambrian basement rocks, and at the eastern and southern margin to palaeogene basalts. minor palaeogene sediments are found at the western and eastern margins. the intrusion was emplaced during flexuring of the continental margin, was tipped to the south after emplacement, and during the late stages of solidification, it was intruded by gabbroic sills and subsequently by alkaline and transitional dyke swarms (nielsen 1978; wotzlaw et al. 2012). the bulk composition of the magma is debated, not only because it is difficult to sum up the compositions of the intrusion, but also because the magma chamber was filled in by multiple pulses of magma over a very short time span (hoover 1989: holness et al. 2007; annen et al. 2022). in addition, only a small part of the roof is preserved, and only assumptions can be made as to what has been lost to erosion. all proposed compositions, however, point to a bulk composition comparable to tholeiitic lavas contemporaneous with the intrusion and present during the opening of the north atlantic. a bulk composition comparable to that of evolved basalt (magnesium number (mg#) c. 45–50) of the geickie plateau formation is suggested by andreasen et al. (2004) and nielsen (2004), whereas jakobsen et al. (2010) and cho et al. (2022), preferred a composition comparable to that of the underlying milne land formation. nielsen et al. (2019b), however, supported a correlation to geikie plateau formation lavas based on comparatively simple, but seemingly robust, major element characteristics. the bulk magma crystallised inward from roof, walls and floor in an onion ring-like structure. the intrusion is subdivided into the layered series (ls) in the floor, the marginal border series (mbs) on the walls, and the upper border series (ubs) below the roof of the magma chamber (fig. 1). the three series are further subdivided into zones and subzones based on parallel evolutions in liquidus parageneses (salmonsen & tegner 2013). in ls, they include the hidden zone (hz; of the lowermost and non-exposed gabbros), lower zone (lz), middle zone (mz) and upper zone (uz). the division is based on the occurrence of liquidus olivine, which is mg-rich in the lz, generally absent in the mz, and fe-rich in the uz. the lz is further subdivided into lza, lzb and lzc l1 l2 l1 smeared inclusion ri ri roof inclusion l2 ml2 ml1.1 ml1.2 ml0 ml2.1 l3 a b figure 2 2 columns fig. 2 the leucogabbro layers (l1, l2 and l3) of the triple group on wagertoppen. a: l1 to l3 on the western face of wagertoppen (1277 m). large sunken blocks of roof gabbro (ri: roof inclusions) disturb the layering. in some cases, the layering drapes around the blocks (photo: m.b. holness). b: l1 and l2 on the nw face of wagertoppen with distinct modal layering of the zebra-banded zone at its base. several sunken roof blocks are fully preserved or smeared (photo: j.c.ø. andersen). macrorhythmic layers ml0 to ml2.1 with leucogabbro tops are indicated by the yellow bars. the upper metres of ml0 host the pd5 mineralisation level, and ml2.1 hosts pd1/au (modified from nielsen et al. 2019a). fig. 3 summary of the skaergaard mineralisation. from left to right: macrorhythmic layers (ml-1 to ml2.2); geochemical subdivisions of the mineralisation, includes the lower pge mineralisation (lpgem), upper pge mineralisation (upgem), upper au mineralisation (uaum) and cu mineralisation (cum, see nielsen et al. 2015 for detailed explanations); density (g/cm3) of bulk rocks and mineralisation levels from centrally-located drill cores; and mineralisation levels (pd6 and pd5 in ml0, pd4a, b in ml1.1, pd3a, b in ml1.2, pd2a, b in ml2 and pd1 and pd1/au in ml2, and au+1 in ml2.1). au+1 (stippled orange line) is only found in drill core 90-18. this concordant mineralisation level is cu-rich and poor in precious metals in all other drill cores. the pge peak in pd5 (designated 0 m in the mineralisation) provides a marker throughout the intrusion (nielsen et al. 2015; fig. 4a). the correlations in the 15–40 m succession of mineralisation levels are based on systematic assays in watts, griffis & mcouat ltd (1991). figure modified from nielsen et al. (2015, 2019a). 3.0 3.1 3.2 3.3 3.4 3.5 density (g/cm³) mineralisation levels 0 40 30 20 10 50 ml2.2 ml2.1 ml1.2 ml1.1 ml-1 ml2 ml0 lp g em u pg em c um u a um pd1 pd5 pd6 pd2 a b pd3a b pd4 a b pd1/au figure 3 1 column au+1 https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 6 of 66 www.geusbul let in.org based on the arrival of clinopyroxene on the liquidus at the base of lzb and, subsequently, titanomagnetite and ilmenite at the base of lzc. the uz is subdivided into uza, uzb and uzc. these subdivisions are based on the arrival of apatite on the liquidus at the base of uzb, and the occurrence of mosaics of grains of ferrohedenbergite replacing inverted β-ferrobustamite at the base of uzc (wager & brown 1968). the same succession in the arrivals of the liquidus phases allows parallel subdivision of the mbs (hoover 1989) into lza*, lzb*, lzc*, mz*, uza* and uzb* (uzc*is not identified) and of the ubs into lza’, lzb’, lzc’, mz’, uza’, uzb’ and uzc’ (salmonsen & tegner 2013; fig. 1). 2.1 the skaergaard pge-au mineralisation the skaergaard pge-au mineralisation, also referred to as the platinova reef, is a complex multi-layered, low-sulphur type deposit found in gabbros rich in feti oxides in the upper 100 m of the mz (bird et al. 1991; andersen et al. 1998). the upper mz is composed of a series of macrorhythmic layers, three of which have distinct leucogabbro tops referred to as l1, l2 and l3. the leucogabbro layers contain up to 70 vol.% plagioclase and are collectively referred to as the triple group. l1, l2 and l3 can be traced by eye across the intrusion (fig.  2). all macrorhythmic layers have comparatively plagioclase-rich tops, and clinopyroxene-rich bases (nielsen et al. 2015). the corresponding more mafic gabbros below the leucogabbro layers are referred to as m1, m2 and m3 (wager & brown 1968). the macrorhythmic layers are commonly separated by a layer with a low concentration of titanomagnetite and ilmenite compared to the rest of the gabbros of the triple group (nielsen et al. 2015). such a 2 m thick layer below m1 is referred to as l0, although it is not to be compared to the plagioclase-rich tops of macrorhythmic layers. in the centre of the intrusion, the precious metals are concentrated in stratiform layers in the lower 50 m of the triple group. marked lateral variations are observed in concentrations and ratios of precious metals, in table 1 specifications of the studied samples sample no. a mineralisation level sample weight (kg) pd pt au pd pt au no. of grains (ppb) (% of (pd+pt+au)) 10-434 pd4a 0.72 529 35 1587 24.6 1.6 73.8 89 10-443 pd5 0.78 2787 137 282 86.9 4.3 8.8 52 10-445 pd5 0.58 2491 129 142 90.2 4.7 5.1 23 18-958 pd1/au 0.97 351 37 283 52.3 5.5 42.2 33 18-972 pd2a 0.80 745 39 114 83.0 4.3 12.7 60 18-978 pd2b 0.80 1051 28 241 79.6 2.1 18.3 157 18-988 pd3b 0.83 987 105 150 79.5 8.5 12.1 33 18-1001 pd4a 1.19 668 74 59 83.4 9.2 7.4 25 18-1010 pd5 0.80 1747 196 126 84.4 9.5 6.1 180 18-1012 pd5 0.72 1240 175 46 84.9 12.0 3.1 116 22-977 au+1 (cu) b 1.05 70 15 140 31.1 6.7 62.2 9 23a-798 pd4a 0.53 360 22 2000 15.1 0.9 84.0 180 23a-806 pd5 0.56 1600 110 250 81.6 5.6 12.8 86 23a-807 pd5 0.76 2800 140 270 87.2 4.4 8.4 165 23a-808 pd5 0.71 2300 120 52 93.0 4.9 2.1 116 24-1018 pd1/au 0.54 100 20 4500 2.2 0.4 97.4 320 24-1022 pd1 0.70 1046 132 814 52.5 6.6 40.9 172 24-1024 pd2-pd1 0.99 30 10 10 60.0 20.0 20.0 4 24-1030 pd2a 1.16 690 20 97 85.5 2.5 12.0 60 24-1034 pd2b 0.92 1300 90 467 70.0 4.8 25.1 271 24-1042 pd3a 1.45 620 20 97 84.1 2.7 13.2 83 24-1045 pd3b 1.02 580 40 52 86.3 6.0 7.7 247 24-1048 pd4a 1.14 630 50 52 86.1 6.8 7.1 122 24-1053 pd4b 0.98 720 20 24 94.2 2.6 3.1 78 24-1056 pd5 0.78 2000 190 112 86.9 8.3 4.9 116 24-1057 pd5 0.78 2800 170 97 91.3 5.5 3.2 373 24-1059 pd5 1.10 1200 110 30 89.6 8.2 2.2 102 24-1062 pd5 1.30 720 110 24 84.3 12.9 2.8 33 bs0304 c pd2 2.33 430 50 2130 16.5 1.9 81.6 490 bs0301 c pd5 1.94 2830 170 240 87.3 5.2 7.4 594 a sample number is comprised of the drill core number (e.g. 24) followed by sample depth (m). for example, sample no. 24-1062 refers to material sampled at 1062–1063 m in drill core 90-24. b cu-rich and au-poor sample from au+1 mineralisation level in drill core 90-22. c bs refers to bulk sample (tof). mineralisation levels (pd5, pd4b, etc.) are named after the dominant precious metal. hs-01 separator was used for sample 23a-806, -807, -808, 18-1010 and 24-1057, hs-02m was used for sample 23a-798 and hs-11 for all others. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 7 of 66 www.geusbul let in.org precious metal mineralogy, parageneses and in concentrations of cu (andersen et al. 1998; nielsen et al. 2005, 2015, 2019a). the central part of the mineralisation is skewed slightly sw of the geographic centre of the intrusion (red star in fig. 1; nielsen et al. 2019a). here, the skaergaard pge-au mineralisation is hosted in macrorhythmic gabbro layers ml0, ml1.1, ml1.2, ml2, ml2.1 and ml2.2 (fig. 3). the macrorhythmic layers were named after the related leucogabbro layers, including l1 and l2 of the triple group. each macrorhythmic layer hosts two mineralisation levels: a lower and an upper (nielsen et al. 2019a; fig. 3; table 1). the pd-rich mineralisation levels are numbered from the top as pd1/au, pd1, pd2a, pd2b, pd3a, pd3b, pd4a, pd4b, pd5 and pd6 (fig. 3). the naming of the mineralisation levels is reminiscent of the early phases of exploration. mineralisation levels, macrorhythmic layers and lithological layers have been shown to maintain near-constant stratigraphic thicknesses and separation across the 0 1000 2000 3000 4000 975 985 995 1005 1015 1025 1035 1045 au m l 2 m l 1 m l 0 pd 2 a a b b a b pd 3 cu pd 6 pd 5 pd 1 l0 l2 l1 m l 2. 1 0 1000 2000 3000 4000 442 452 462 472 482 492 502 512 l0 l2 l1 pd 2 pd 3 pd 4 pd 5 a a a b b b (b) 0 2000 4000 6000 749 759 769 779 789 799 809 819 l2 l1 l0 0 1000 2000 3000 4000 –12 –2 8 18 28 38 48 58 l2 l1 l0 3.0 3.2 3.4 3.6 975 985 995 1005 1015 1025 1035 1045 l2 l1 l0 l2 l1 l0 749 759 769 779 789 799 809 819 3.0 3.2 3.4 0 1000 2000 3000 4000 5000 6000 7000 –600 –400 –200 0 200 400 90-14 90-15 90-13 90-20 90-22 90-24 90-19 90-23a 90-25a w e top of l3 pd1 pd5 a b density (g/cm³) density (g/cm³)pd+pt and au (ppb), cu (ppm) pd+pt and au (ppb) pd+pt and au (ppb), cu (ppm) pd+pt and au (ppb) d ep th in c or e or c hi pl in e (m ) 90-22 core or chip line 90-22 90-17a 90-23a middag chip line 90-23a distance from centre (m) 289 289 2150 3440 4700 3440 distance from western margin (m) figure 4 full page, portrait fig. 4 correlation between elemental concentrations and lithology of host gabbro in the skaergaard pge-au mineralisation. a: pge, au and cu concentrations and density profiles of the skaergaard pge-au mineralisation. the correlations demonstrate the continuity of layering, mineralisation levels and lateral variations in elemental distributions. blue: pge (pd+pt); yellow: au, and red: cu. from left to right: density profile and (pd+pt), au and cu in the centrally located drill core 90-22, 25 cm bulk-rock samples; (pd+pt) and au in drill core 90-17a located 1150 m from the western margin, 1 m bulk-rock samples; (pd+pt), au and cu in drill core 90-23a located 900 m from the eastern margin of the intrusion; and (pd+pt), au and density profile in the middag chip line profile, c. 5 km n of the central mineralisation (locations in fig. 1). grey shading indicates the elevation of leucogabbro layers l0, l1 and l2 of the triple group (see nielsen et al. 2015, 2019a for further information). modified from nielsen et al. 2019a. b: cross-section of the skaergaard intrusion (core locations in fig. 1b; for locations of cores not used in this study see nielsen et al. 2019a) with projected mineralisation intersects (to scale). elevations of pd5 (blue) and pd1 (red) mineralisation levels and leucogabbro layer l3 (yellow) are relative to sea level. the profile only includes data from drill cores pd5, pd1 and l3. mineralisation levels and the lithological layering are concordant across the 7000 m wide and >600 m deep bowl-shaped succession of mz gabbros. slightly modified from nielsen et al. (2019a). fig. 5 visualisation of the 3d structure of the precious metal–rich part of the mineralisation (not to scale). the stack of pge-rich mineralisation levels with upward-decreasing diameter within the bowl-shaped succession of macrorhythmic layers. blue shading: pge-rich; green shading: au-rich; yellow shading: au-rich but pge-poor layers. the maximum concentration of au is in the uppermost mineralisation level in the central part of the intrusion. modified from nielsen et al. (2019a). not to scale figure 5 1 column https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 8 of 66 www.geusbul let in.org intrusion (fig. 4a). the stratigraphically lowest mineralisation levels, pd6 and pd5, are hosted in macrorhythmic layer ml0. pd5 is the main pge resource of the deposit. it is 5 m thick at a cutoff c. 0.7 g/t (ppm) in the central parts of the intrusion where it has an average grade of c. 2 g/t (pge+au) over 5 m. pd5 can be up to 8 m thick at the same cutoff grades near the margin of the intrusions (nielsen et al. 2019a). ml1.1 hosts mineralisation levels pd4b and pd4a, ml1.2 hosts mineralisation levels pd3b and pd3a, ml2 hosts mineralisation levels pd2a and b, ml2.1 hosts mineralisation level pd1 and pd1/au (fig. 3; nielsen et al. 2019a), ml2.2 hosts the pge-poor mineralisation level au+1. in general, the au/pge ratio increases up through the mineralisation levels (andersen et al. 1998; nielsen et al. 2005; holwell & keays 2014) from c. 0.028 in pd6 (drill core 90-22; bernstein & nielsen 2005) to a maximum of 13 in the au+1 peak in drill core 90-18 (watts, griffis & mcouat ltd. 1991). the upward increase in au/pge is also seen in individual mineralisation levels due to fractionation in mushes of the macrorhythmic layers. above the uppermost au-rich mineralisation level follows a series of at least eight stratiform and cu-rich levels with negligible concentrations of pge and au (nielsen et al. 2015). in the most northerly exposures of ml0, concentrations of pge and au overlap within 1 m of ml0 gabbro (turner & mosher 1989; andersen et al. 1998). in drill cores from near the margins of the intrusion, pge is concentrated only in the mineralisation levels pd5 and pd6 of ml0, whereas au is concentrated in pd4a and b of ml1.1. all mineralisation levels above pd4a are rich in precious metals in their geographical centres whereas their marginal zones are rich in cu and poor in pge and au. the width of the cu-rich marginal zone increases up the mineralisation levels. the transition from au-rich and cu-poor (<100 ppm) uppermost precious metal mineralisation level to the overlying level rich in cu (>1000 ppm) climbs stepwise up mineralisation levels and macrorhythmic layers, and close in on the geographical centre of the mineralisation (star in fig. 1). the distribution of the precious metals can be visualised as a stack of gold-rimmed plates with an upwardly-decreasing diameter (figs 4b, 5) separated by gabbro that is poor in precious metals and cu. the stratigraphic separation between the main pd horizon, pd5, and the uppermost au layer pd1/au, therefore, increases from <1 m at the margins to >43 m in more central parts of the intrusion. 2.2 the bowl shape of the triple group contrary to general perceptions, the macrorhythmic layers, the leucogabbro layers of the triple group (l1– l3 plus l0), as well as the mineralisation levels are not planar, nor did they form initially as horizontal layers in the intrusion; they are all bowl-shaped. wager & deer (1939) proposed a bowl shape for the ls succession, which was subsequently reconstructed by nielsen (2004) and nielsen et al. (2015, 2019a). the bowl shape depicted in fig. 4b was constructed from the elevation relative to sea level of the tops of l3, pd5 and pd1 mineralisation levels projected on to a profile across the intrusion (nielsen et al. 2019a). details for this type of reconstruction are found in nielsen et al. (2015, 2019a) and holness et al. (2017). the three markers are concordant and maintain near constant separation, confirming that mineralisation levels and the lithological layering are concordant (nielsen et al. 2015). instead of a stack of horizontal plates, the spatial distribution of the precious metals is best compared to a stack of goldrimmed bowls of upward-decreasing diameter (fig. 5). the distribution of precious metals in five consecutive macrorhythmic layers in the centre of the mineralisation and the concentration of the precious metals in just one macrorhythmic layer near the margins of the intrusion does not reflect telescoping of mineralisation layers. it demonstrates significant differences in the mobility of precious metals and other metals such as cu in mush melt at the margins and near the geographical centre of the mineralisation (rudashevsky et al. 2014, 2015; nielsen et al. 2015, 2019a). https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 9 of 66 www.geusbul let in.org 3 samples and methods the precious metal mineralogy of the skaergaard pge-au mineralisation was studied in detail from a total of 30 samples (table 1). the data presented here were extracted from laboratory reports (nielsen et al. 2003a–e; cabri et al. 2005a; rudashevsky & rudashevsky 2005a, b, 2006a, b; rudashevsky et al. 2009a,  b, 2010a–d, 2012а–i). subsequent publications such as rudashevsky et al. (2014, 2015), nielsen et al. (2005, 2015, 2019a) and godel et al. (2014), used data from these reports, and include descriptions of new minerals (rudashevsky et al. 2004; mcdonald et al. 2008 and paragenetic information (nielsen et al. 2019a). the studied samples stem from five profiles covering top to base of the precious metal mineralisation and bulk samples collected for metallurgical tests. samples were collected from drill core 90-23a near the eastern margin, three central profiles from drill cores 90-18, 90-24 and the bulk sampling site referred to as ‘toe of forbindelsesgletscher’ (tof; samples bs0301 and bs0304), and drill core 90-10 from near the western margin (table 1; locations in fig. 1). in addition, we include one sample from drill core 90-22 to compensate for a stratigraphic interval in drill core 90-24 lost to basaltic dykes. drill core 90-22 is collared approx. 350 m sw of drill core 90-24 and is the geochemical mirror image. samples from tof and drill cores 90-18, 90-22 and 90-24 are all considered central, based on their geographical positions >1 km in from the margin of the intrusion. all sample preparation and analytical methods are described in detail in rudashevsky et al. (2001, 2002, 2014, 2015). here, we briefly outline the methods used in the preparation and analysis of the samples presented in this study. one metre sections of drill core (size bq, 36.5 mm diameter) were cut along their length. the material used in the study represents approx. one third of the circular cross-section of the drill core. the weight of each sample was between 0.5 and 1.3 kg, depending on previous use of drill core material for exploration purposes and recovery tests. table 1 gives the sample identification data including the drill core number, the 1 m depth interval that was sampled, the weight of the submitted sample, the assay for the sample interval, and the relative proportions of pd, pt and au, as well as the number of precious metal grains recovered. samples were carefully crushed and sieved to preserve primary grain sizes and shapes, and subsequently subjected to separation using hydroseparator cnt hs-01, cnt hs-02m and cnt hs-11 (rudashevsky & rudashevsky 2006c, 2007). the heavy mineral concentrates of the samples were divided into three grain-size fractions, 125–80 µm, 80–40 µm and <40 µm, and studied under the microscope. the grains were imaged and analysed using a camscan 4dv scanning electron microscope (sem) equipped with a link an-10000 spectrometer for electron microprobe (emp) analyses. the methods are described in detail in, for example, rudashevsky et al. (2018). the size of individual grains is approximated as the area of the circle that encompasses the grain (i.e. the equivalent circle diameter, ecd). elemental concentrations are given as wt%, ppm or ppb. sample numbers and scales are embedded in individual backscattered electron images (bse; i.e. in figs 6–9, 15–20 and 23–32). individual grains of rock-forming minerals, sulphides and precious metal phases are identified by the unique fig. 6 backscattered electron (bse) images of the petrographic relations between rock-forming silicates, oxides, sulphides and pge phases in polished thin sections (panels a–c, e, g–l) and polished monolayer thin sections of hs concentrates (panels d, f). a–g: samples from the central part of the intrusion (drill cores 90-18, 90-24 and tof). h–l: samples from more marginal parts of the intrusion (drill cores 90-10 and 90-23a). abbreviated names of minerals, compounds and alloys are as follows: bn and bn: bornite. ch and cc: chalcocite. cpx and cpx: clinopyroxene. opx and opx: orthopyroxene. pyr (exs): сpx-opx exsolution. pl and pl: plagioclase (includes pl i with an38–42; and pl ii with an82). timt: titaniferous magnetite. ilm: ilmenite. ol: olivine. mt: magnetite. hb: hornblende. act: actinolite. bt: biotite. q: quartz. srp: serpentine. tlc and tlc: talc. chl: chlorite. sk and sk: skaergaardite. (cu,pd) α: pd-cu alloy. apd: arsenopalladinite. grain numbers and scales included in the images. table 3 provides standard formulae for the precious metal minerals. formulae for common rock-forming minerals and base metal sulphides are found in appendix 1. a b c d e f g h i j k l fig. 6 2 columns 24-1022, ps p.l 300 μm bs0301, p.5 30 μm 24-1057, p.15 30 μm 18-978, 80 #1 30 μm 1048, ps p.6 100 μm 24-1045, 125 #14 30 μm 24-1057, p11 300 μm 10-445, p.s. #4 30 μm 10-434, ps #5 100 μm 23a-808, p.s.808-2 p.4 100 μm 23a-808, p.s.808-1 p.2 100 μm 23a-808 p.s. 808-2 p.2 100 μm https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 10 of 66 www.geusbul let in.org number of the drill core from which concentrates were made, the start depth of the selected 1 m interval, the grain-size fraction and the number (#) of the individual grain. for example, 23a-806, 45/2 #2 is the unique identification of grain #2 in mount 45/2 from the interval 806 to 807 m in drill core 90-23a. a total of >4000 grains were studied. in addition, the petrographic relations of alloys and minerals dominated by cu and pd to hosts and sulphide minerals were imaged in situ by godel et al. (2014) using high resolution x-ray computed tomography (hrxct) on small, drilled cores samples (8 and 4 mm in diameter). samples were collected from the pge-rich lower mineralisation level (pd5) from tof and the upper au-rich mineralisation level (pd4) near the margin on kraemer ø. this latter sample is equivalent to the au-rich samples 10-434 and 23a-798 included in the present study. see appendix 1 for a list of the names and formulae of the common rock-forming minerals, rare accessory phases and common base metal sulphides in the studied samples. the precious metals mentioned in the text and their standard formulae are listed in table 3. fig. 6 (continued) backscattered electron (bse) images of the petrographic relations between rock-forming silicates, oxides, sulphides and pge phases in polished thin sections (panels a–c, e, g–l) and polished monolayer thin sections of hs concentrates (panels d, f). a–g: samples from the central part of the intrusion (drill cores 90-18, 90-24 and tof). h–l: samples from more marginal parts of the intrusion (drill cores 90-10 and 90-23a). abbreviated names of minerals, compounds and alloys are as follows: bn and bn: bornite. ch and cc: chalcocite. cpx and cpx: clinopyroxene. opx and opx: orthopyroxene. pyr (exs): сpx-opx exsolution. pl and pl: plagioclase (includes pl i with an38–42; and pl ii with an82). timt: titaniferous magnetite. ilm: ilmenite. ol: olivine. mt: magnetite. hb: hornblende. act: actinolite. bt: biotite. q: quartz. srp: serpentine. tlc and tlc: talc. chl: chlorite. sk and sk: skaergaardite. (cu,pd)α: pd-cu alloy. apd: arsenopalladinite. grain numbers and scales included in the images. table 3 provides standard formulae for the precious metal minerals. formulae for common rock-forming minerals and base metal sulphides are found in appendix 1. a b c d e f g h i j k l fig. 6 2 columns 24-1022, ps p.l 300 μm bs0301, p.5 30 μm 24-1057, p.15 30 μm 18-978, 80 #1 30 μm 1048, ps p.6 100 μm 24-1045, 125 #14 30 μm 24-1057, p11 300 μm 10-445, p.s. #4 30 μm 10-434, ps #5 100 μm 23a-808, p.s.808-2 p.4 100 μm 23a-808, p.s.808-1 p.2 100 μm 23a-808 p.s. 808-2 p.2 100 μm https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 11 of 66 www.geusbul let in.org fig. 7 bse images of sulphides. panels a–e are polished thin sections of gabbroic host. panels f–n are polished monolayer thin sections of heavy mineral hs concentrates. images include samples from the central part of the intrusion (a–e, g–n), and one sample from the marginal part of the intrusion (f). note exsolution textures in panels k–n. abbreviated names of minerals, compounds and alloys are as follows: bn and bn: bornite. ch and cc: chalcocite. cp: chalcopyrite. dgn: digenite. cpx: clinopyroxene. opx: orthopyroxene. pl: plagioclase. ol: olivine. timt: titaniferous magnetite. ilm: ilmenite. ap: apatite. bt: biotite. chl: chlorite. grain numbers and scales included in the images. table 3 provides formulae for the precious metal minerals. formulae for common rock-forming minerals and base metal sulphides are found in appendix 1. a b c d e f g h i j k l m n 24-1024, 125 #3b 300 μm 1048, ps p.1 100 μm 24-1057, p.6 30 μm 24-1022, ps p.1 30 μm bs0304, p.2 30 μm 23a-808, 75 #s12 30 μm bs0304-125 30 μm 10-445, 40 #1b 10 μm 18-1012, 40 #42a 10 μm 24-1053, 63 #4b 10 μm bs0301-125-2 30 μm 24-1057, 75 glb9 30 μm 24-1057, 125 glb11 30 μm 18-1010, 125-2 #sb 30 μm https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 12 of 66 www.geusbul let in.org figure continues.... fig. 8 2 columns a b c d e f g h i j k 24-1056, 125 #6 30 μm bs0301-45-2-1 #1 10 μm 24-1053, 63 #14 10 μm 1048, 40 #31 10 μm bs0301-125-1 #11 30 μm 24-1018, 80 #5 10 μm bs0301-75, 2 #28 10 μm 23a-806, 125 #3 30 μm 23a-807, 45 #1 30 μm 24-1030, 125 #8d 30 μm 23a-806, 75 10 μm fig. 8 bse images of rare minerals in polished monolayer thin sections. table 3 provides formulae for the precious metal minerals. formulae for common rock-forming minerals and base metal sulphides are found in appendix 1. formulae are given here for rare minerals that are not listed in table 3 or appendix 1. abbreviated names of minerals, compounds and alloys are as follows: copn and pn-(co): cobaltian pentlandite (panels a–d). pn: pentlandite (panels e–f). po: hexagonal pyrrhotite (panel g). gn: galena (panels h–j). sph: sphalerite (panel k). zns: sphalerite (panel l). cds: greenockite/ hawleyite (panels m, n). cls: clausthalite (pbse) (panel o). aspy: arsenopyrite (panel p). spn: thiospinel group, carrollite (cu(co,fe,ni)2s4) (panel q). atc: atacamite (cu2cl(oh)3) (panel r). sk: skaergaardite (panels a, b, e, g, l). (cu,pd): cu-pd alloy (panel c). nls: nielsenite (pdcu3) (panel c). at: atokite (pd3sn) (panel e). aucu: tetra-auricupride (panels f, q). (au,ag): native gold (panel h). pd3ag2s: coldwellite (panel n). vsl: vasilite ((pd,cu)16s7) (panel r). bn: bornite. ch: chalcocite. cp: chalcopyrite (panels h, o). cpx: clinopyroxene (panel j). opx: orthopyroxene (panels j, q). pl: plagioclase (panel q). ilm: ilmenite (panels i, p). sk: skaergaardite (panel l). grain numbers and scales included in the images. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 13 of 66 www.geusbul let in.org fig. 8 (continued) 2 columns l m n o p q r 24-1057, 45-1 #4 10 μm 18-988, 40 #2a 10 μm 18-988, 40 #22 10 μm 10-434, 70 #1a 30 μm bs0301-45 10 μm bs0304 75 #29 30 μm bs0301 75-1 #28 30 μm cls cp fig. 8 (continued) bse images of rare minerals in polished monolayer thin sections. table 3 provides formulae for the precious metal minerals. formulae for common rock-forming minerals and base metal sulphides are found in appendix 1. formulae are given here for rare minerals that are not listed in table 3 or appendix 1. abbreviated names of minerals, compounds and alloys are as follows: copn and pn-(co): cobaltian pentlandite (panels a–d). pn: pentlandite (panels e–f). po: hexagonal pyrrhotite (panel g). gn: galena (panels h–j). sph: sphalerite (panel k). zns: sphalerite (panel l). cds: greenockite/hawleyite (panels m, n). cls: clausthalite (pbse) (panel o). aspy: arsenopyrite (panel p). spn: thiospinel group, carrollite (cu(co,fe,ni)2s4) (panel q). atc: atacamite (cu2cl(oh)3) (panel r). sk: skaergaardite (panels a, b, e, g, l). (cu,pd): cu-pd alloy (panel c). nls: nielsenite (pdcu3) (panel c). at: atokite (pd3sn) (panel e). aucu: tetra-auricupride (panels f, q). (au,ag): native gold (panel h). pd3ag2s: coldwellite (panel n). vsl: vasilite ((pd,cu)16s7) (panel r). bn: bornite. ch: chalcocite. cp: chalcopyrite (panels h, o). cpx: clinopyroxene (panel j). opx: orthopyroxene (panels j, q). pl: plagioclase (panel q). ilm: ilmenite (panels i, p). sk: skaergaardite (panel l). grain numbers and scales included in the images. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 14 of 66 www.geusbul let in.org a b c d e f bs0304 125 #s28 30 μm 24-1018, 63 #35a 30 μm 24-1018, 80 #1a 30 μm 24-1018, 63 #1a 30 μm 24-1018, 40 #62 10 μm 24-1018, 63 #20 10 μm fig. 9 bse images of sulphides from the au-rich pd1/au horizon in the central part of the intrusion. polished monolayer thin sections of the heavy mineral hs concentrates. table 3 provides formulae for the precious metal minerals. formulae for common rock-forming minerals and base metal sulphides are found in appendix 1. formulae are given here for rare minerals that are not listed in table 3 and appendix 1. abbreviated names of minerals, compounds and alloys are as follows: bn: bornite. ch: chalcocite. mln: melonite ((ni,pd)te2). te: native tellurium. orc: orcelite (ni5-xas2). un1: unnamed compound (cu,fe)3tes2. ilm: ilmenite. chl: chlorite. grain numbers and scales included in the images. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 15 of 66 www.geusbul let in.org 4 mineralogy, composition and petrographic relationships 4.1 the gabbroic host the petrography, mineral chemistry and bulk geochemistry of the gabbros of the skaergaard intrusion are described in numerous papers and monographs (e.g. wager & brown 1968; mcbirney 1996). in the triple group and in the mineralisation levels, the liquidus phases show limited compositional variation with cores of plagioclase crystal (~an46) and clinopyroxene (mg# ≈ 58). the mineralised gabbros, however, show wider compositional ranges in cores of plagioclase grains (an40–52), exsolved clinopyroxene (mg# = 59–65) and orthopyroxene (mg# = 43–56). the gabbros of the skaergaard pge-au mineralisation are complex rocks that include two main silicates and feti oxide parageneses (nielsen et al. 2015). a liquidus paragenesis of plagioclase-pyroxene and feti oxides, an interstitial paragenesis of plagioclase-pyroxene-olivine-feti oxides and apatite related to the sulphide and precious metal mineralisation, and a late paragenesis with hydrous minerals formed in reaction with fluids (see section 5.4). in the inner parts of the intrusion, the host gabbros contain irregular aggregates of feti oxides that are separated from earlier resorbed plagioclase and pyroxenes by zones of fayalitic olivine (mg# = 40–56) and ca-rich plagioclase ii (up to an89; fig. 6a–g). following nielsen et al. (2015), the olivine-rich rims are developed at the interface between the first-formed liquidus paragenesis and a later interstitial parageneses dominated by feti oxides (fig. 6a–g). this reaction relationship is not seen close to the margin of the intrusion. here, the space between cumulus phases is commonly filled by aggregates of, for example, serpentine and magnetite (fig. 6h), and no reaction is seen at contacts between feti oxides and clinopyroxene (fig. 6i). in samples from drill core 90-23a near the eastern margin, the primary minerals of the gabbros are variably followed by hydrous silicates such as biotite, hornblende, actinolite, chlorite, ferrosaponite, ferrous talc, serpentine, quartz, magnetite and calcite (fig. 6j–l). these lateto post-magmatic minerals sometimes form veins in pristine host gabbros (fig. 6k). apatite, baddeleyite, zircon, thorianite, monazite and uraninite are examples of accessory minerals identified by emp analysis in the polished thin sections of the studied bulk-rock samples as well as in polished monolayer samples and mounts prepared from concentrates obtained by hydro-separation (hs) concentrates (see methods, chapter 3). 4.2 the ore minerals 4.2.1 the sulphides the gabbros enriched in precious metals always have small proportions of cu-fe sulphides, c. 0.5  vol.% and the bulk-rock samples mostly have <80 ppm cu (nielsen et al. 2015). bornite is the dominant sulphide, followed by chalcocite digenite and chalcopyrite. the grain size is rarely more than a few tenths of µm and rarely exceeds 0.1 mm. details of the petrography, parageneses and compositions of the sulphides in the samples studied here are found in nielsen et al. (2003a–e), rudashevsky et al. (2004, 2009a, b, 2010a–d, 2012a–i, 2014, 2015), cabri et al. (2005a), rudashevsky & rudashevsky (2005a, b, 2006a, b) and mcdonald et al. (2008). the shapes and habits of sulphide grains include three general types as follows: type 1: irregularly shaped aggregates from the interstices between rock-forming minerals (e.g. figs 6b, c, 7c) type 2: sphere-shaped grains with a sphericity up to 0.95 and droplet-like micro-globules (figs 6d, f, 7a, b, d, e, g–n) type 3: aggregates of small flaky grains in association with lateto post-magmatic minerals. most common are type 1 sulphides grains, located between the rock-forming minerals of the gabbroic host (figs 6b, c, 7c). the droplet-like inclusions (type 2) are comparatively rare but occur throughout the mineralisation in ilmenite (figs 6d, f, 7a), titaniferous magnetite (figs 6g, 7b), pyroxene grains (figs 6a, 7d) and plagioclase and olivine (figs 6c, 7e). type 3 sulphide grains are limited to samples from near the margins of the intrusion in drill cores 90-10 and 90-23a, where the sulphides often occur in intergrowths with late-magmatic hydrous silicates (figs 6j–l, 7f). the sulphide grains can be monomineralic and composed of bornite chalcocite, digenite or chalcopyrite (fig. 7g–j) or be composed of two or three different sulphides in variable volumetric proportions. polymineralic cu-fe sulphide grains typically exhibit exsolution textures, for example, of intermediate solid solution (iss; fig. 7k–n). the cu-fe sulphide paragenesis varies from centre to margin of the intrusion. samples from the central parts of the intrusion (tof and drill cores 90-18, 90-22 and 90-24; fig. 1) have comparably s-poor paragenesis, with chalcocite and digenite and intergrowths of bornite and chalcocite. the s-poor paragenesis is suggested to have exsolved from a parental phase with a composition https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 16 of 66 www.geusbul let in.org close to digenite (e.g. andersen 2006). samples from near the margins of the intrusion (drill cores 90-23a and 90-10) have a more s-rich sulphide association dominated by bornite grains and aggregates of bornite and chalcopyrite (nielsen et al. 2003a, d, e; rudashevsky et al. 2012c–e). digenite is also present, but only in subordinate amounts. in addition to the dominant cu-fe sulphides, polished thin sections and heavy mineral hs concentrates reveal a suite of rare sulphides (table 2). they form small inclusions, 1–20 µm in size, in cu-fe sulphide grains, in rock-forming minerals and inside grains of precious metal minerals. these sulphides include cobaltian pentlandite (co,ni,fe)9s8, pentlandite (ni,fe,co)9s8, galena (pbs), hexagonal pyrrhotite (fe7s8), sphalerite (zn,fe)s, greenockite/hawleyite (cds), in addition to phases such as clausthalite (pbse; fig. 8) and carrollite (cu(co,fe,ni)2s4). we note that the petrographic relations suggest that cobaltian pentlandite with up to 7.8 wt% rhodium (rh) and 5.5 wt% pd, formed contemporaneously with the pge minerals (table 2, rows 2 and 3). the sulphide parageneses of the uppermost au-rich mineralisation in a given drill core is very different, irrespective of the stratigraphic position in the stack of macrorhythmic layers and mineralisation levels. the sulphide parageneses of au-rich samples are rich in tellurium (te) with inclusions of melonite (ni,pd)te2, native te and orcelite (ni5-xas2), which is found as inclusions in tetra-auricupride (aucu; fig. 9f; table 2, row 11), as well as unnamed sulpho-tellurides of cu and fe. the sulpho-tellurides have stoichiometric compositions and several grains are observed. three types of sulpho-tellurides have the empirical formulae: (cu,fe)5tes2, (cu,fe)3tes2 and (cu,fe)6tes2 (fig. 9a–e; table 2, rows 7–10), and are believed to represent unnamed minerals. we believe the grain textures show that such sulpho-tellurides formed by re-equilibration and exsolution (fig. 9c) from an unknown metastable, higher-t magmatic phase. table 2 composition of rare sulphides, sulpho-tellurides, tellurides and arsenides row no. sample no. grain no. cu fe ni co zn pd rh s as te sum cobaltpentlandite (co,ni,fe)9s8 1 multiple ave. of 9 grains wt% 0.9 12.5 19.4 32.4 n.d. 0.6 0.8 32.6 n.d. n.d. 99.1 prop. 0.11 1.76 2.63 4.35 0.04 0.06 8.04 17.00 sd 1.03 3.73 4.14 6.17 0.60 0.80 0.41 2 24-1024 80#2 wt% 3.3 14.9 14.9 29.1 n.d. 5.5 n.d. 32.2 n.d. n.d. 99.9 prop. 0.42 2.14 2.03 3.96 0.41 8.04 17.00 3 18-1010 75/2#5 wt% 1.5 6.8 17.3 34.4 n.d. n.d. 7.8 31.8 n.d. n.d. 99.6 prop. 0.20 0.99 2.40 4.74 0.61 8.06 17.00 pentlandite (ni,fe,co)9s8 4 multiple ave. of 12 grains wt% 1.0 17.8 30.7 17.2 n.d. n.d. n.d. 32.8 n.d. n.d. 99.5 prop. 0.12 2.49 4.09 2.24 8.01 17.00 sd 1.39 6.41 4.08 5.04 carrollite cu(co,fe,ni)2s4 5 bs0304 125#3 wt% 21.6 3.3 3.5 30.8 n.d. n.d. n.d. 40.0 n.d. 0.8 100.0 prop. 1.06 0.19 0.19 1.64 3.91 0.02 7.00 sphalerite (zn,fe)s 6 multiple ave. of 14 grains wt% 0.2 1.1 n.d. n.d. 65.7 n.d. n.d. 32.9 n.d. n.d. 99.9 prop. n.d. 0.02 0.98 1.00 2.00 melonite (ni,pd)te2 7 24-1018 40#62 wt% n.d. n.d. 69.3 n.d. n.d. 2.0 n.d. n.d. n.d. 29.7 100.0 prop. 0.97 0.03 2.00 3.00 unnamed (cu,fe)5tes2 8 24-1018 & bs0304 ave. of 27 grains wt% 55.4 5.8 n.d. n.d. n.d. n.d. n.d. 11.9 n.d. 26.4 99.5 prop. 4.49 0.54 1.91 1.07 8.00 sd 2.23 1.34 0.62 0.83 unnamed (cu,fe)3s2te 9 bs0304 ave. of 7 grains wt%. 40.6 7.4 n.d. n.d. n.d. n.d. n.d. 17.0 n.d. 14.7 99.7 prop. 2.44 0.50 2.02 1.04 6.00 sd 2.13 0.67 1.03 1.90 unnamed cu6s2te or (cu,fe)2(s,te) 10 24-1018 ave. of 4 grains wt% 63.4 0.9 n.d. n.d. n.d. n.d. n.d. 11.1 n.d. 23.9 99.3 prop. 5.80 0.10 2.02 1.09 9.00 sd 1.19 0.52 0.30 0.79 orcelite ni5-xas2 11 24-1018 63#24 wt% n.d. n.d. 69.3 n.d. n.d. n.d. n.d. n.d. 29.7 n.d. 100.0 prop. 4.90 3.10 8.00 ave.: average. prop.: atomic proportions. sd: standard deviation. n.d.: not detected. indicates no data. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 17 of 66 www.geusbul let in.org table 3 precious metal minerals of the skaergaard pge-au mineralisation row no. mineral formula no. of grains total area (∑, μm2) no of samples d ecd (μm) f total area (%) e min. ave. max. intermetallic compounds and alloys dominated by pd and cu 1 skaergaardite a pdcu 2054 1124677 24 57 1 23 93 2 nielsenite b pdcu3 73 21081 11 1.07 2 15 50 3 (pd,cu) alloy (pd,cu,au) 13 1644 6 0.08 3 8 20 4 (cu,pd)β (cu,pd,au,pt) 135 36891 5 1.87 1 19 51 5 (cu,pd)α (cu,pd) 65 22600 2 0.87 2 16 36 6 alloy (cu,pt) (cu,pt) 2 125 1 0.01 2 7 11 7 native pd pd 2 936 2 0.05 20 29 29 intermetallic compounds and alloys dominated by au, ag and cu 8 tetra-auricupride aucu 1006 378984 20 19.2 2 19 57 9 auricupride cu3au 27 5953 4 0.30 3 16 31 10 unnamed au3cu 64 14885 17 0.75 2 14 35 11 alloy (au,cu,pd) 156 56815 17 2.88 2 17 54 12 alloy (cu,au) 25 11256 5 0.57 7 15 41 13 alloy (au,cu,ag,pd) 284 114772 8 5.81 4 15 46 14 alloy (au,cu,fe,pd) 2 912 1 0.05 11 21 31 15 alloy (cu,au,ni,zn) 1 620 1 0.03 28 16 native au c (au,ag) 15 3958 7 0.20 1 11 40 17 alloy (ag,cu) 6 4964 3 0.25 14 29 58 18 native ag (ag,au) 9 2456 6 0.12 5 16 38 intermetallic compounds and alloys with pt 19 alloy (pt,cu,fe,pd) 21 634 7 0.03 1 5 14 20 alloy (pt,fe,cu,pd) 18 680 9 0.03 1 6 11 21 alloy (pt,cr) 1 2345 1 0.12 55 sulphides with precious metals 22 vasilite c (pd,cu)16s7 119 32281 21 1.64 1 19 58 23 vysotskite (pd,ni,cu)s 148 32708 14 1.66 1 18 45 24 unnamed (сu,pd)2s 5 253 1 0.01 2 7 12 25 unnamed (pd,cu,pt)3s2 12 3495 4 0.18 2 14 33 26 coldwellite pd3ag2s 7 429 4 0.02 4 8 20 27 unnamed pd3(ag,cd,cu,tl)s 1 80 1 0 10 28 unnamed (pd,hg,ag)2s 1 54 1 0 6 29 argentite/acanthite ag2s 1 1786 1 0.09 48 30 stephanite ag5sbs4 1 3710 1 0.19 69 31 polybasite (ag,cu)16sb2s11 1 605 1 0.03 28 32 unnamed pd3cu6(te,sn)2s3 1 48 1 0 8 arsenides with precious metals 33 vincentite pd3as 92 18161 8 0.92 1 13 59 34 arsenopalladinite/ pd8as3 90 19921 6 1.01 3 16 61 stillwaterite c 35 palladoarsenide pd2as 7 1101 3 0.06 4 14 25 36 isomertieite pd11as2sb2 7 399 3 0.02 1 8 20 37 sperrylite ptas2 53 1420 7 0.07 1 6 69 38 majakite pdnias 2 18 1 0 3 4 4 stannides and tellurides with precious metals 39 unnamed pd11as2sn2 1 100 1 0.01 11 40 keithconnite c pd3-x(te,pb,sn) 160 9319 25 0.47 1 9 25 41 kotulskite pdte 29 1301 7 0.07 2 7 13 42 merenskyite pdte2 1 85 1 0 10 43 unnamed pd2te 2 310 1 0.02 12 15 18 44 sopcheite ag4pd3te4 1 48 1 0 8 45 telargpalite (pd,ag)3(te,pb) 1 45 1 0 8 46 hessite ag2te 1 74 1 0 10 47 unnamed (pd,ag)2te 4 66 1 0 3 5 7 48 unnamed pd2cutebi 1 49 1 0 8 49 zvyagintsevite c pd3(pb,te,sn) 279 30837 22 1.56 1 13 69 50 atokite c pd3(sn,te,pb) 68 4218 12 0.21 1 8 24 51 cabriite pd2cusn 7 188 3 0.01 4 6 21 52 alloys (pd,cu,sn,te,pb) 44 3581 14 0.18 1 7 23 53 naldrettite pd2(sb,sn) 1 10 1 0 10 bismuth and pd minerals 54 froodite pdbi2 1 180 1 0.01 18 total     5129 1974138   100       a accepted by ima (rudashevsky et al. 2004). b accepted by ima (macdonald et al. 2008). c minerals recorded prior to use of hydroseparation techniques (nielsen et al. 2005). d number of samples in which the mineral was identified. e area of phase compared to the total area of precious metal mineral grains in each sample. f average equivalent circle diameter (ecd) of precious metal mineral grains. indicates no data. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 18 of 66 www.geusbul let in.org concentrations of chalcophile elements such as pd, au, ag, te and se are recorded in the cu-fe sulphide blebs from pd5 and pd4–pd2 mineralisation levels in drill core 10-48 (holwell et al. 2015). elements such as te (50–100 ppm) and se (300–500 ppm) only show minor variations, whereas elements such as pd, au and pt vary by several orders of magnitude. variations in ratios between the precious metals and cu can, however, be due to several processes, and conclusions are difficult to reach. based on petrographic observations and bulk-rock compositions, nielsen et al. (2015) argued for syn-magmatic dissolution of primary immiscible sulphide droplets leading to enrichment in mainly pd and subsequently to late au-enrichment caused by mobility and migration of au in mush melts and fluids. some textures are apparently the result of immiscibility between cu-s melts and pd-rich melts (nielsen et al. 2015) and could account for melt droplets composed of >50 vol.% skaergaardite (pdcu) and with an extreme concentration of pd. some sulphide droplets host large crystals of pd minerals. others show necking, which is understood to be the separation between precious metal crystals and sulphide melt due to crystallisation of neighbouring phases or gravitational and due to the difference in density between the precious metal grains and the related sulphide melt. such sulphide melts would be depleted in precious metals. some sulphide blebs, however, have extremely high contents of precious metals (e.g. holwell et al. 2015). they may be caused by tiny particles of precious metal mineral below the surface of the studied grains and would lead to strongly elevated and non-representative precious metal concentrations in the sulphide hosts. holwell et al. (2015) also reported especially high concentrations of se in sulphide blebs from the upper au-rich mineralisation level pd1/au in the centre of the intrusion and in agreement with bulk-rock enrichment in se reported by keays & tegner (2016). 4.2.2 precious metal minerals only nine precious metal phases were identified and named in the skaergaard pge-au mineralisation prior to our use of hydrosepation to produce heavy mineral concentrates. andersen et al. (1998) found (cu,fe)(au,pd,pt) alloy with variable au/pge ratios to be the totally dominant precious metal phase together with small amounts of electrum (au,ag), atokite (pd3sn), zvyagintsevite (pd3pb), vasilite ((pd,cu)16s7), keithconnite (pd3+x(te,as)), melonite ((ni,pd,cu,fe)(te,sb)2), arsenopalladinite (pd8(as,sb,sn)3) and unnamed (pd,cu)2s (bird et al. 1991). table 3 lists the precious metal minerals identified in the 30 samples, together with information on grain sizes (ecd) and the calculated relative abundances. the phases are subdivided into the following groups or associations: 1. intermetallic compounds (commonly stoichiometric) and alloys dominated by pd and cu with minor substitutions of pt, fe, cu and cr. the group represents 61.3% of the entire paragenesis (table 3, rows 1–7, 19–21) 2. intermetallic compounds and alloys of au, ag and cu representing 30.1% of the parageneses (table 3, rows 8–18) 3. sulphides of pd and cu with substitutions of ag, cd, hg and tl, representing 3.8% of the paragenesis (table 3, rows 22–34) 4. arsenides of pd with pt and ni substitutions representing 2.1% (table 3, rows 35–38) 5. stannides and tellurides with precious metals including intermetallic compounds of pd and cu with sn, pb, te, (sb, bi) representing 2.7% (table 3, rows 39–54, including froodite (pdbi2)). all the original data can be found in nielsen et al. (2003a–e), rudashevsky & rudashevsky (2005a, b, 2006a, b), rudashevsky et al. (2009a, b, 2010a–d, 2012а– i, 2014, 2015) and cabri et al. (2005a). in total, 54 different compositions are now reported, three of which were approved as new minerals by the international mineralogical association (ima). they include skaergaardite (pdcu; rudashevsky et al. 2004), nielsenite (pdcu3; mcdonald et al. 2008) and naldrettite (pd2sb; cabri et al. 2005b, rudashevsky & rudashevsky 2006b). the dominant pd-mineral skaergaardite has the simplified formula pdcu, but is better described as (pd,au,pt) (cu,fe,zn,sb,sn,pb,te) and the dominant au-mineral tetra-auricupride (aucu). the formula for the prevalent (cu, fe)(au,pd,pt) alloy of andersen et al. (1998) covers both of these phases. all other minerals are of minor importance and include zvyagintsevite (pd3pb; 5%), keithconnite (pd3-xte; 3%), vasilite (pd16s7; 2%), arsenopalladinite (pd8(as,sb)3; 2%) and native gold (0.3%; table 3). these are averages for the entire mineralisation and are not representative of any given sample. the very large number of individual grains in the study (tables 1, 3) ensures statistically robust information for the approximate volumetric proportions of the precious metals throughout the intrusion, the morphology of individual grains, the paragenetic relations and the compositional variations. all of these are used for the reconstruction of primary parageneses (nielsen et al. 2019a). here, we summarise the mineralogical data and focus on the genesis of the main levels of the mineralisation, including the lower pge-rich level across the intrusion, and the upper au-rich mineralisation level as recorded in bulk sample bs0304 and drill cores 90-18 https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 19 of 66 www.geusbul let in.org and 90-24 near the geographic centre of the mineralisation, 90-23a from the eastern margin and 90-10 from the western margin of the intrusion (fig. 10). figure 10 illustrates the relative proportions of phases in the main pd mineralisation level (pd5) and the au-rich upper mineralisation levels. the data shown are averages based on all the samples up though pd5 and the sections through the au-rich level in a given drill core. the full data set is found in appendix 6 of nielsen et al. (2019a). skaergaardite (pdcu) is the dominant pd-cu precious metal mineral of the main pge mineralisation level (pd5) in drill core 90-24 from near the geographic centre of the mineralisation (fig. 10). au-bearing phases do not exceed 2 vol.%. south-west of the centre, in drill core 90-18, pd5 has an elevated proportion of precious metal sulphides and arsenides (fig. 10). but near the margins on either side of the intrusions the pd-cu phases are almost absent (<2%; “cu-pd” in fig. 10), despite comparatively high pge concentrations and an up to 8 m thick pd5 at a cut off of 0.7 ppm (watts, griffis & mcouat ltd. 1991). at the western margin in drill core 90-10, pd5 is dominated by 47.2% palladium sulphides followed by 26.3% arsenides and 9.1% intermetallic compounds of pb, sn, te and summing to 82.6%. the remaining c. 17% of the paragenesis is composed of au-bearing minerals (fig.  10). in stark contrast, pd5 at the eastern margin (drill core 90-23a) is dominated by intermetallic compounds of pb, sn and te (51%) including zvyagintsevite (pd3pb), atokite (pd3sn), keithconnite (pd3-xte), and arsenides (25.6%; fig. 10) including arsenopalladinite (pd8as3), vincentite ((pd,pt)3(as,sb,te); in early reports referred to as guanglinite), palladoarsenide (pd2as), and sulphides (7.8%; fig. 10), including vasilite (pd16s7) and vysotskite  (pds); total 84.4%. au-bearing phases and cu-pd account for c. 15% of the precious metal paragenesis (fig. 10). in the central drill core 90-24, the au-rich precious metal paragenesis of the pd1 and pd1/au mineralisation levels is totally dominated by au-cu minerals and phases mainly in the form of tetra-auricupride (aucu). the pd1–pd1/au interval in drill core 90-24 is 2 m thick at a cutoff at 1 ppm au and has average au 4.2 ppm (data in watts, griffis & mcouat ltd. 1991). the pd1–p1/au interval in drill core 90-18 is dominated by pd-cu phases including skaergaardite (pdcu). the pd1/au mineralization level has only 28.8 % au-cu phases compared to >90% in nearby drill core 90-24 ( fig.  10). most gold is in drill core 90-18 concentrated 14 m above pd1 in the au+1 mineralisation level (nielsen et al. 2015, 2019a) where the au phases constitute 94 vol.% of the precious metal paragenesis. immediately above au+1 follows mineralisation levels that are strongly elevated in cu (>1000 ppm) and very poor in precious metals (nielsen et al. 2015, 2019a). at the margins in drill cores 90-10 and 90-23a, the au-rich pd4a mineralisation level is located 10 m above pd5 (fig. 10). at a cut-off of 0.7 ppm, the assays for the margin centre margin west east sul fig. 10 2 columns int 1.3% 1.6% sul 1.3% int au-cu 28.8 % int 0.8% 11.8% ars 25.6%ars 26.3% int 9.1% sul 47.2% sul 1.1% au-cu 17% (cu-pd) 0.4% au-cu 14.1 % cu-pd 1.5% sul 7.8% int 51% au-cu 78.2%au-cu 92.6% int 3.1%int 0.8% ars 18.7% ars 5.5% ars 0.3% au-cu 4.4% au-cu l.7% cu-pd 95.4% cu-pd 82.2% cu-pd 70.4% au-cu 98.5% int 0.2%1.3% sul pg eri ch a uri ch core 90-10 (pd4) core 90-18 (pd1/au)* core 90-24 (pd1/au) core 90-23a (pd4) core 90-10 (pd5) core 90-18 (pd5) core 90-24 (pd5) core 90-23a (pd5) fig. 10 contrasts between precious metal parageneses across the intrusion in the upper au-rich mineralisation and the lower pge-rich mineralisation level (pd5). the data are summarised for drill cores 90-10 near the western margin, 90-18 from the sw centre, 90-24 from the centre, and 90-23a near the eastern margin (drill core locations in fig. 1; original data in electronic appendix 12 of nielsen et al. 2019a). the lower pge-rich levels are perfectly concordant with variations in pd/pt ratios up the layered gabbros despite the stark contrasts in parageneses. the au-rich mineralisation levels are found increasingly higher in the succession of gabbro layers towards the centre of the intrusion (see also fig. 4). the au-rich mineralisation levels are pd1/au in core 90-24, pd1/au* and in au+1 in 90-18 and pd4a in cores 90-10 and 90-23a. total areas of precious metals are recalculated to 100% and used as approximations to volume relations. cu-pd: minerals, intermetallic compounds and alloys of cu and pd that may be substituted by pt, fe, cu and/or cr; au-cu: minerals, intermetallic compounds and alloys of au and cu (ag); sul: sulphides of pd, cu (ag, cd, hg, tl); ars: arsenides of pd (pt, ni); int: intermetallic compounds of pd, cu, sn, pb, te (со, sb, bi). https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 20 of 66 www.geusbul let in.org pd4a level show 2.4 ppm au over a width of 0.8 m in drill core 90-10 and 0.6 m in drill core 90-23a. 4.2.3 intermetallic compounds and alloys of cu and pd intermetallic compounds and alloys of cu and pd are the main precious metal minerals in the skaergaard pge-au mineralisation. they range from native palladium to cu-pd alloys with variable proportions or trace amounts of other elements (rudashevsky et al. 2015). they include the intermetallic compounds skaergaardite (pdcu; rudashevsky et al. 2004) and nielsenite (pdcu3; mcdonald et al. 2008). traditional studies of polished thin sections of gabbro from pd5 revealed only a few skaergaardite grains (fig. 6a–e, g). whereas heavy mineral hs concentrates of the same samples have provided information on composition, shape and paragenesis of more than 2340 grains of cu-pd compounds and alloys from 26 of the 30 studied samples (table 3). the compositional range in the intermetallic compounds of cu and pge in all samples and mineralisation levels is shown in fig. 11. the subdivision of the suite of minerals and phases is based on the atomic proportion (at.%), calculated as the % of (pd+pt+au) in cu + (pd+pt+au). the average compositions of the identified phases in the suite from cu to pd are listed in table 4. the suite (fig. 11) divides into: (1) cu-rich alloy (cu,pd)α, 3–20 at.% (pd+pt+au); (2) nielsenite (pdcu3), 20–30 at.%; (3) (cu,pd)β alloy, 35–45 at.%; (4) skaergaardite (pdcu), 45–55 at.%; (5) solid solution of (pd,au,cu) alloys, 70–75 at.%; (6) solid solution of (pd,cu,au,pt) alloys, 80–90 at.%; (7) native pd, c. 100 at.%. in total, these intermetallic compounds and cu and pd alloys account for 61% of the total area of all the examined precious metal mineral grains from the skaergaard pge-au mineralisation (table 3). skaergaardite (pdcu) dominates (57%), followed by nielsenite (cu3pd; 1%) and non-stoichiometric alloys of cu and pd (0.3–1.9%), and rare grains of (pd,cu,au,pt) alloy and native palladium. the paired correlation coefficients calculated from all available analyses of skaergaardite (pdcu; table 5) show that pd is substituted by au and pt, and that cu is substituted by a wide range of elements including fe, sb, sn, te and pb. the maximum substitutions of these elements in skaergaardite and nielsenite are listed in tables 4 and 6. these observations can be compared to a factor analysis of 311 skaergaardite grains identified in hs heavy mineral concentrates of sample 24-1057, covering the pge peak of level pd5 (fig. 12; rudashevsky et al. 2004). this analysis divides the skaergaardite compositions into four groups characterised by the following elements: factor 1: pd, fe and zn; factor 2: cu, au and te; factor 3: sn and pb; and factor 4: pt. the range in elemental substitutions in 1482 skaergaardite grains from all studied drill core samples is visualised in fig. 13. the ‘pd, fe and zn’ group, represents more than 80% of all analyses and are pd-skaergaardites with low au and pt (table 6, rows 1, 4, 5). in this group, cu is substituted by up to 8.9 wt% fe and 9.5 wt% zn (table 4). the second group includes skaergaardite grains rich in au and te (table 6, rows 3, 7) with typical isomorphic replacement of pd by up to 31.5 wt% au and up to 4.5 wt% te (rudashevsky et al. 2004). they are usually poor in fe and zn. the sn and pb group includes skaergaardite with up to 21.4 wt% sn, and 5.5 wt% pb substitution for cu, fe and zn (rudashevsky et al. 2004 and table 6, rows 6 and 8). these skaergaardites are usually poor in cu, fe and zn. the skaergaardites of the pt group are pt-rich (table 6, row 2) with up to 12.7 wt% pt substituting for pd (rudashevsky et al. 2004). minerals composed of pd and cu dominate in all mineralisation levels (pd6–pd1) in the central drill cores and below the uppermost au-rich mineralisation level n um be r o f a na ly se s 0 20 40 60 80 100 120 140 0–5 5–1 0 10–1 5 15–2 0 20–2 5 25–30 30–35 35–40 40–45 45–50 50–55 55–60 60–65 65–70 70–75 75–80 85–90 80–85 90–95 95–1 00 (cu,pd)α (pd,au,cu) pd (cu,pd) 444 pd+pt+au in cu+pd+pt+au (at.%) nls (pd,cu,au,pt) sk 1036 n = 1680 fig. 11 1.5 column β fig. 11 intermetallic compounds and alloys ranging from cu to pd. based on emp analyses of a total of 1680 grains. minerals, compounds and alloys are identified on the basis of precious metal to cu ratio expressed as (pd+au+pt) in at.%. from left to right: (cu,pd)α alloy: 3–20 at.%; nls: nielsenite (pdcu3), 20–30 at.%; (cu,pd)β alloy: 35–45 at.%; sk: skaergaardite (pdcu), 45–55 at.%; (pd,au,cu) solid solution: 70–75 at.%; (pd,cu,au,pt) solid solution: 80–90 at.%; pd: native pd, c. 100 at.%. note that bars for “sk” are shortened to fit the diagram. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 21 of 66 www.geusbul let in.org (pd3) at bulk-sampling locality tof (location in fig. 1). skaergaardite (pdcu) from the lowermost mineralisation levels (pd5 and pd4) in cores from the centre of the intrusion is significantly enriched in pt and low in au and te, whereas skaergaardite of the upper mineralisation levels (pd2 and pd1) is enriched in au (table 7) and mirrors the bulk-rock concentrations of pge and au (e.g. nielsen et al. 2015). 4.3 relationships between host rocks, sulphides and precious metal phases the petrographic relations between all the phases in the studied samples provide a framework for the interpretation of the order in which the phases have formed. parageneses of coexisting phases can be defined and provide the basis for a combined mineralogical and petrogenetic mineralisation model. table 4 average compositions of cu and pd minerals no. of grains   pd pt au cu fe zn sn te pb sum skaergaardite (pdcu) 1480 ave. (wt%) 58.2 1.6 2.6 30.6 3.8 1.2 0.5 0.4 0.4 99.4 prop. 0.96 0.01 0.02 0.84 0.12 0.03 0.01 0.01 0.00 2.00 sd 5.60 3.37 5.08 2.93 1.67 1.13 1.86 0.69 0.86 max. 36.7 41.9 n.d. 8.9 9.5 21.4 7.2 6.6 nielsenite (cu3pd) 60 ave. (wt%) 31.0 2.7 3.0 62.1 0.6 n.d. n.d. n.d. n.d. 99.4 prop. 0.89 0.04 0.05 2.99 0.03 4.00 sd 5.20 4.80 5.80 3.90 0.50 max. 25.6 25.9 n.d. 2.8 1.4 1.5 1.9 (cu,pd)β alloy 67 ave. (wt%) 53.9 0.5 0.9 41.3 1.5 0.7 0.1 0.2 0.5 99.5 prop. 0.42 0.00 0.01 0.54 0.02 0.01 0.00 0.00 1.00 sd 2.30 1.20 1.30 2.80 1.10 0.90 0.30 0.40 0.80 max. 6.0 7.9 n.d. 6.7 5.1 1.9 2.0 2.6 (cu,pd,pt)α alloy 64 ave. (wt%) 19.4 1.3 0.3 77.8 0.7 n.d. n.d. n.d. 0.1 99.4 prop. 0.13 0.00 0.00 0.86 0.01 1.00 sd 6.18 2.82 0.61 5.75 0.40 0.31 max. 16.3 2.6 n.d. 1.7 1.5 (pd,au,cu) alloy 5 ave. (wt%) 41.7 1.0 40.6 11.6 1.7 n.d. 1.0 0.9 n.d. 98.5 prop. 0.47 0.01 0.25 0.22 0.04 0.01 0.01 1.00 sd 16.49 0.78 20.04 0.51 0.69 1.90 1.85 max. 1.9 54.0 n.d. 2.5 3.8 3.7 (pd,cu,au,pt) alloy 4 ave. (wt%) 53.3 29.2 9.1 5.6 1.8 n.d. n.d. n.d. n.d. 99.0 prop. 0.49 0.10 0.15 0.21 0.06 1.00 sd 17.97 20.66 10.42 4.05 0.98 max. 76.8 55.9 24.6 11.3 3.1 ave.: average. prop.: atomic proportions. sd: standard deviation. max.: maximum content (wt%) of element in single grain of the given mineral. n.d.: not detected. indicates no data. table 5 correlation coefficients for elements contained in 1482 grains of skaergaardite element pt au cu fe zn sn te pb pd –0.53 –0.81 –0.13 +0.34 +0.35 –0.33 pt   –0.28 +0.24 –0.21 au   –0.41 –0.24 +0.38 cu   –0.41 –0.25 –0.23 fe   –0.36 –0.2 –0.36 zn   –0.21 –0.15 –0.13 sn   –0.14 +0.21 te               only correlations with p ≥ 0.95 are shown. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 22 of 66 www.geusbul let in.org 4.3.1 droplets and globules droplets of sulphide melt in liquidus feti oxides are conspicuous in the skaergaard pge-au mineralisation (nielsen et al. 2005; godel et al. 2014). the droplets formed from sulphide melt, and some contain crystals and globules of precious metal phases. the globules are rounded grains and may have uncertain origins (see e.g. karup-møller et al. 2008). droplets of sulphide melt may also be trapped between crystallising rock-forming minerals to form anhedral grains. many skaergaardite grains in hs concentrates are associated with cu-fe sulphide grains, which may or may not represent droplets deformed between the grains of rock-forming minerals (e.g. fig. 6b, c; see also holwell et al. 2015). sulphide droplets with micro-droplets of skaergaardite and (pd,cu) alloys are found across the intrusion in ilmenite and titaniferous magnetite of the pd5 and pd4 mineralisation levels (fig. 6d–g). a droplet of skaergaardite without a visible relationship to sulphide phases was also observed (fig. 6e). feti oxide grains also contain spherical and ‘octahedral’ silicate inclusions, composed of hornblende, biotite, olivine, ferrosaponite (fig. 14a–c) and individual (liberated) sulphide droplets with hydrous silicates (fig. 14d–g) and skaergaardite globules (fig. 14h–k). in some sulphide droplets, skaergaardite accounts for most of their volume (e.g. fig. 14j). table 6 composition of varieties of skaergaardite and nielsenite a row no. variety pd pt au cu fe zn sn te pb sum skaergaardite 1 pd wt% 63.6 n.d. n.d. 29.3 5.0 2.6 n.d. n.d. n.d. 100.5 prop. 1.00 0.78 0.15 0.07 2.00 2 pt wt% 33.4 36.7 n.d. 20.5 8.6 0.6 n.d. n.d. n.d. 99.8 prop. 0.63 0.38 0.66 0.31 0.02 2.00 3 au wt% 25.7 n.d. 41.9 29.7 0.7 n.d. n.d. 1.0 n.d. 99.0 prop. 0.51 0.46 1.00 0.02 2.00 4 fe wt% 63.7 n.d. n.d. 26.9 8.9 n.d. n.d. n.d. n.d. 99.5 prop. 1.01 n.d. 0.72 0.27 2.00 5 zn wt% 62.4 n.d. n.d. 25.1 1.7 9.5 n.d. n.d. n.d. 98.7 prop. 1.01 0.69 0.05 0.25 2.00 6 sn wt% 56.7 n.d. n.d. 18.8 2.0 n.d. 20.5 n.d. 1.1 99.1 prop. 1.02 0.57 0.07 0.33 0.01 2.00 7 te wt% 29.5 9.3 24.6 24.5 2.9 n.d. n.d. 7.4 1.6 99.8 prop. 0.57 0.10 0.26 0.81 0.11 0.12 0.03 2.00 8 pb wt% 57.1 3.9 0.0 30.6 0.9 0.8 0.9 0.0 6.4 100.6 prop. 0.97 0.04 0.00 0.87 0.03 0.02 0.01 0.00 0.06 2.00 nielsenite 9 pd wt% 35.8 0.0 0.0 62.7 1.3 n.d. n.d. n.d. n.d. 99.8 prop. 1.00 0.00 0.00 2.93 0.07 4.00 10 pt wt% 15.8 25.6 1.1 54.6 2.8 n.d. n.d. n.d. n.d. 99.9 prop. 0.50 0.44 0.02 2.87 0.17 4.00 11 au wt% 18.6 1.8 25.9 52.0 1.3 n.d. n.d. n.d. n.d. 99.6 prop. 0.60 0.03 0.45 2.84 0.08 4.00 12 au-pt wt% 18.1 9.5 15.0 55.5 1.2 n.d. n.d. n.d. n.d. 99.3 prop. 0.57 0.16 0.26 2.94 0.07 4.00 a maximum substitutions of indicated element in grain of given mineral. prop.: atomic proportions. n.d.: not detected. indicates no data. atomic proportions normalised to 2.00 for skaergaardite and to 4.00 for nielsenite. pd pb pb pd pd sn sn sn a b zn zn zn pt pt pt cu cu cu te au te au te auiv fe fe fe i i i ii ii ii iii iii iii pt iv iv –0.8 –0.5 –0.2 0.1 0.4 0.7 1 f2 factor 2 sn 0 .8 1, p b 0. 70 fe 0 .7 3, c u 0. 26 , z n 0. 23 –0.9 –0.6 –0.3 0 0.3 0.6 0.9 f1 au 0.88, te 0.73, cu 0.55 pd 0.87, fe 0.53, zn 0.42 f1 au 0.88, te 0.73, cu 0.55 pd 0.87, fe 0.53, zn 0.42 –0.9 –0.6 –0.3 0 0.3 0.6 0.9 –0.6 –0.3 –0.8 –0.5 –0.2 –0.1 0.4 0.7 1.0 0 0.3 0.6 0.9 f3 fa ct or 3 pt 0 .8 7, s n 0. 35 , f e 0. 26 zn 0 .5 0, c u 0. 40 , p d 0. 34 –0.6 –0.3 0 0.3 0.6 0.9 –0.9 –0.6 –0.3 0 factor 1 0.3 0.6 0.9 fig. 12 2 column c fig. 12 factor analysis on the skaergaardite (pdcu) composition from sample 90-24-1057 in level pd5. a total of 311 analyses are included in the modelling. a: f1 (factor 1) versus f2 (factor 2). b: f1 (factor 1) versus f3 (factor 3). c: 3d diagram of all three factors. the compositions indicated by the stippled lines include group i: (pd-fe-zn); group ii: (cu-au-te); group iii: (pb-sn); group iv: pt-rich skaergaardite. modified from rudashevsky et al. (2004). https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 23 of 66 www.geusbul let in.org godel et al. (2014) conducted a hrxct investigation of two samples from the mineralisation. the recomposed figures show silicate inclusions and sulphide droplets inside ilmenite grains (fig. 15a, b), and precious metal grains as trains of elongated inclusions next to feti oxide (fig. 15c). the relative volume of sulphide and hydrous silicate inclusions to precious metal grains is lower in the trains, compared to the volume relations in liquidus feti oxides (godel et al. 2014). in one sample from the au-rich uppermost mineralisation level from pd4 near the western margin of the intrusion (equivalent to sample 10-434), godel et al. (2014) found silicate droplets in feti oxide hosts, no sulphide droplets, but grains of au phases along boundaries of rock-forming minerals. 4.3.2 cu-pd minerals 4.3.2.1 relationships between cu-pd minerals and cu-fe sulphides skaergaardite (pdcu), nielsenite (cu3pd), non-stoichiometric alloys (cu,pd)β and (cu,pd)α compounds are mostly located at the margin of sulphide grains, irrespective of 0 au pt sn <1 1–5 5–10 10–15 15–20 20–25 25–30 30–35 35–40 200 400 600 800 1000 1200 # a # = 1482 40–45 wt% 0 10 20 30 40 50 60 70 80 % <1 1–5 5–10 10–15 15–20 20–25 25–30 30–35 35–40 40–45 wt% b au pt sn 0 200 400 600 800 1000 1200 # <1 1–2 2–3 3–4 4–5 5–6 6–7 7–8 wt% fe zn te pb c 0 5 10 15 20 25 30 35 40 % <1 1–2 2–3 3–4 4–5 5–6 6–7 7–8 wt% fe zn te pb d e 0 10 20 30 40 50 60 70 80 % <1 1–2 2–3 3–4 4–5 5–6 6–7 7–8 wt% figure 13 2 columns fig. 13 histograms illustrating frequency of substitutions in skaergaardite based on a total of 1482 emp analyses. a: au, pt, sn substitutions vs. number of analyses (#). b: au, pt, sn substitutions vs. frequency (f) in %. c: fe, zn, te, pb substitutions vs. number of analyses. d: fe and zn substitutions vs. frequency. e: te and pb substitutions vs. frequency. table 7 average composition of skaergaardite in mineralisation levels pd5–pd4 and pd2–pd1 no. of grains   pd pt au cu fe zn sn te pb sum pd mineralisation levels: pd5 and pd4 ave. (wt%) 57.9 2.3 2.5 29.9 4.1 1.3 0.6 0.4 0.4 99.3 829 prop. 0.95 0.02 0.02 0.82 0.13 0.04 0.01 0.01 0.00 sd 5.26 3.87 4.69 2.13 1.75 1.06 1.94 0.69 0.88 pd-au mineralisation levels: pd2 and pd1 ave. (wt%) 53.8 0.6 8.3 30.6 3.0 1.0 0.9 0.6 0.3 99.1 56 prop. 0.91 0.01 0.08 0.87 0.10 0.03 0.02 0.01 0.00   sd 9.17 1.85 9.39 1.48 1.36 0.78 1.48 1.09 0.63 ave.: average. prop.: atomic proportions. sd: standard deviation.indicates no data. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 24 of 66 www.geusbul let in.org their shape. however, numerous droplet-like sulphide globules also host well-shaped micro-globules composed of cu-pd compounds. they also dominate the precious metal paragenesis in anhedral, interstitial intergrowths with cu-fe sulphides. the volume ratios between host sulphide and precious metal compounds show large variations (figs 16, 17). the spatial relationship between sulphide host and micro-droplets include: (1) rounded droplets or partly faceted grains of сu-pd compounds at the margin of sulphide globules (figs 16a–e, 17b, e, h, k) (2) small specks of precious metal droplets in sulphide host (figs 16g, h, 17a, c, f) (3) microdroplets of sulphides inside larger precious metal droplets (figs 16f, 17d) (4) fine precious metal emulsion in sulphide host (figs 16i, j, 17a) (5) small precious metal droplets merged into larger grains of skaergaardite (figs 16h–l, 17d). euhedral crystals of precious metal phases are also found in some sulphide grains or droplets. they may be cubic or cubo-octahedral (figs 16m–p, 17l). the relationships described here between sulphide hosts and precious metal phases apply to the entire range of cu-pd compounds including (cu,pd)β alloy, skaergaardite (pdcu), nielsenite (cu3pd) and (cu,pd)α alloy. the heavy mineral hs concentrates contain fully liberated (free) grains of cu-pd compounds, including numerous droplet-shaped and irregularly shaped grains. they may be composed of one or more grains, and some resemble amalgamated micronuggets (fig.  18). small amounts fig. 14 2 columns a b c d e f g h i j k bs0304 p.11 100 μm 18-1010, 125-1 #3r 30 μm 18-1010 10 μm 24-1057, 125-1 #12a 30 μm bs0301-125-1 30 μm 18-972, 80 #8 30 μm 24-1022, 80 #8 30 μm 24-1057, 75 #24 30 μm bs0301 75-2 #46 10 μm 24-1057, 45-1 #66 10 μm 24-1022, 80 #7 30 μm fig. 14 bse images of silicate globules in ilmenite and examples of hydrous silicates coexisting with sulphides and skaergaardite in polished thin sections and polished monolayer thin sections of gabbroic host and in heavy mineral hs concentrates. table 3 provides formulae for the precious metal minerals. formulae for common rock-forming minerals and base metal sulphides are found in appendix 1. abbreviated names of minerals, compounds and alloys are as follows: ilm: ilmenite. cpx: clinopyroxene. opx: orthopyroxene. ol: olivine. hb: hornblende. bt: biotite. act: actinolite. fspn: ferrosaponite. chl: chlorite. ct and ct: calcite. bn: bornite. ch: chalcocite. sk: skaergaardite. kth: keithconnite. pl: plagioclase. grain numbers and scales included in the images. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 25 of 66 www.geusbul let in.org of bornite or chalcocite attached to these ‘free’ precious metal globules suggest an origin with sulphide droplets or anhedral sulphide grains from interstitial spaces. the 2054 grains of skaergaardite (pdcu) found in the hs concentrates are 1–93 μm in size, with an average of 23 μm (table 3). godel et al. (2014) found a very similar average of 19.2 μm for 25 skaergaardite grains by hrxct. 4.3.2.2 intergrowths of more cu-pd minerals more complex grains composed of two coexisting precious metal compounds, with or without cu-fe sulphide, are also found. the observed precious metal compound pairs include: skaergaardite (sk1, pdcu; sk2, au-rich (pd,au) cu; fig. 16c), skaergaardite-nielsenite (fig. 17g), alloy (cu,pd)β-nielsenite (fig. 17h), nielsenite-alloy (cu,pd)α (fig. 17m). the two types of skaergaardite (sk1 and sk2; fig. 16c) have 9.0 and 23.3 wt% au in solid solution and the two types of (cu,pd)α alloy in figure 17m have 16.4% and 27.4% pd, respectively (rudashevsky & rudashevsky 2006a). 4.3.2.3 relationship between cu-pd minerals and other precious metal phases cu-pd minerals hosted in sulphides are often accompanied by other precious metal phases. they include alloys a c e f d b b 1 mm 100 μm 25 μm 500 μm 500 μm 100 μ m cusulfide pd-cu alloy inclusion spherical inclusion octahedral inclusion feti oxide feti oxide inclusion cusulfide cusulfide pd-cu alloy pd-cu alloy fig. 14 2 columns fig. 15 recompiled high-resolution x-ray computed tomography (hrxct) images after godel et al. (2014) showing 3d images of the relationship between feti oxides (grey), cu-, and fe-sulphides (yellow), skaergaardite (red) and silicate inclusions (blue) enclosed in feti oxides. a: view into an ilmenite grain with silicate inclusions and sulphide droplet with skaergaardite grain attached. b: magnified image of the sulphide droplet. c: trains of silicate and sulphide with skaergaardite, possibly along pathway for interstitial melt. d: focussed view shows the high skaergaardite to sulphide ratio compared to the ratio for sulphide droplets enclosed in ilmenite. e: focussed view of the perfectly spherical inclusions with hydrous silicates and the octahedral silicate inclusion rich in feti oxides. the latter, also referred to as negative crystals, formed from silicate melt that was trapped while the ilmenite host crystallised. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 26 of 66 www.geusbul let in.org of pt, fe, cu and pd; intermetallic compounds of pd, pb, sn, te, sb; pd-sulphides, and au-cu-ag minerals (fig. 19). they are spatially related to and often hosted in cu-pd minerals and alloys, and form a single, but complex precious metal paragenesis. whereas the cu-pd minerals skaergaardite (pdcu) and (cu,pd) alloy, as a rule, are found in the interior of sulphide grains, other precious metal phases are mostly located at the margins of their cu-pd hosts. these include the dominant au and ag minerals ch+bn sk bn fig. 16 2 columns a b c d e f g h i j k l m n o p bs0301 125-2 #9 30 μm bs0301 45 #180 10 μm 24-1057, 45-1 #162 10 μm 24-1057, 45-1 #154 10 μm 24-1034, 40 #33 10 μm bs0301 45 #173 10 μm 24-1057, 45-1 #16 10 μm 18-1010, 125-2 #3 30 μm 24-1057, 75 #68 30 μm 24-1034, 40 #30 10 μm bs0301 45 #99 10 μm 18-978, 63 #17 10 μm 24-1057, 75 #21 30 μm 18-978, 80 #1 10 μm 18-978, 63 #17 10 μm 24-1053, 40 #39 10 μm fig. 16 bse images showing the petrographic relations between skaergaardite and cu-fe sulphides in polished monolayer thin sections of hs concentrates. table 3 provides formulae for the precious metal minerals. formulae for common rock-forming minerals and base metal sulphides are found in appendix 1. note the multitude of relations from crystals of skaergaardite in centres as well as at margins of sulphide droplets, to a wide range of exsolution textures, and to tiny, disseminated droplets of skaergaardite throughout complex cu-fe sulphide grains. abbreviated names of minerals, compounds and alloys are as follows: sk: skaergaardite including sk1: (pdcu)) and sk2: (pd,au)cu). bn: bornite; ch: chalcocite. cp: chalcopyrite. sample numbers and scales included in the images. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 27 of 66 www.geusbul let in.org tetra-auricupride (aucu; fig. 19i) and unnamed au3cu (fig. 19k), native silver (fig. 19l), and varieties of precious metal minerals (fig. 19b–h) including varied and often poly-mineral intergrowths of pd-sulphides like vasilite (pd16s7) and vysotskite (pds), intermetallic compounds with sn, pb and te including the minerals zvyagintsevite (pd3pb), atokite ((pd,pt)3sn), keithconnite (pd3-xte) and cabriite (pd2cusn). some may have exsolved from the cooling cu-pd host. we note that the factor analysis of skaergaardite (pdcu) compositions in rudashevsky et al. (2015) and observed petrographic relations summarised in this study show that the cu-pd phases in intergrowths with pt minerals are usually pt-rich, that skaergaardite in intergrowths with au and te minerals have significant au and te substitution, and that skaergaardite intergrowths with sn intermetallic compounds are sn-rich. there is a clear relationship between the composition of the cu-pd host and the spatially related precious metal paragenesis. 4.3.2.4 pt minerals the gabbros of the mineralisation are in general poor in pt with averages over 1 m of drill core of <0.1 ppm but reach a maximum of 0.25 ppm in the pd5 peak (watts, griffis & mcouat ltd. 1991). pt minerals are exceedingly rare and account for <0.2% of all precious metal mineral grains in all studied samples (table 3, rows 19–21). they are usually <10 µm in size and occur as inclusions in cu-fe sulphide grains, pd-rich pges and cu-au minerals. pt is hosted in the following: 1. (pt,fe,pd,cu) alloys in nine of the investigated samples (fig. 20b, c, e; table 8, rows 1 and 2) 2. (pt,cu,fe,pd) alloy in a total of 21 grains found in 7 samples (fig. 20a, d; table 8, row 2) 3. 53 grains of sperrylite (ptas2) from 7 samples (fig. 20h–l) 4. two grains of (cu,pt) alloy from a single sample (fig. 18g; table 8, row 3) fig. 17 2 columns a b c d e f g h i j k l m 24-1034, 40 #80 10 μm 24-1045, 40 #18 10 μm 18-1012, 40 #26 10 μm 24-1045, 40 #26 10 μm 24-1034, 80 #4 30 μm 18-1012, 40 #85 10 μm 24-1057, 75-22 #1 10 μm 24-1045, 40 #92 10 μm 1045n 123 #10 30 μm 24-1045, 40 #146 10 μm 24-1045, 40 #9 10 μm 24-1045, 63 #8 10 μm 24-1045, 125 #15 30 μm fig. 17 bse images showing the petrographic relationships between nielsenite (cu3pd), (pd,cu) alloys, skaergaardite and cu-fe sulphides in polished sections of hs concentrates. note the multitude of petrographic relations and relative proportions between phases. abbreviated names of minerals, compounds and alloys are as follows: sk and sk: skaergaardite (pdcu). nls and nls: nielsenite (pdcu3). (cu,pd), (cu,pd,pt)β, (cu,pd)α, (cu,pd)β: alloys of cu, pd and pt. ch and cc: chalcocite. bn and bn: bornite. table 3 provides formulae for the precious metal minerals. formulae for common rock-forming minerals and base metal sulphides are found in appendix 1. grain numbers and scales included in the images. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 28 of 66 www.geusbul let in.org 5. one grain of an exotic (pt,cr,pd) alloy (fig. 18f; table 8, row 4) the pt alloys are mainly found in the pd5 mineralisation level in more central parts of the mineralisation ( fig. 20a, b). in drill cores from the margins of the intrusion (90-10 and 90-23a), pt is hosted in sperrylite (ptas2) associated with pd arsenides, pge intermetallic compounds and au minerals (fig. 20c–e, i–l). the largest grain of sperrylite has an ecd of 69 μm. 4.3.3 intermetallic compounds and alloys of au, cu and ag minerals composed only of au and cu are rare in nature (e.g. novgorodovа 1983; bird et al. 1991; andersen et al. 1998; spiridonov 2010a). only two are presently approved by the ima. these are tetra-auricupride (aucu) and auricupride (aucu3). tetra-aurucupride is tetragonal and is the structural analogue of artificial aucu i (chen et al. 1982; novgorodova 1983; spiridonov et al. 2005). auricupride is an ordered cubic phase and the structural analogue of synthetic aucu3 i (okamoto et al. 1987). to these are related more disordered cubic alloys (solid solutions) of au and cu (e.g. lozhechkin 1935, 1939; novgorodova 1983; spiridonov et al. 2005). zaccarini et al. (2004) report the occurrence of cu-rich tetra-auricupride with a simplified composition close to cu3au2 and intermediate between tetra-auricupride (aucu) and unnamed (aucu3). grains with the stoichiometric composition au3cu are common in the skaergaard intrusion, but this unnamed mineral with cubic syngony is at present insufficiently documented for ima approval (pokrovskii et al. 1979; spiridonov & pletnev 2002; knight & leitch 2001). the cu-au (ag) intermetallic compounds and alloys in the studied samples exhibit a wide range of compositions. they form a compositional suite from auricupride (aucu3), to tetra-auricupride (aucu), unnamed au3cu, and native gold with variable substitutions by ag. all the minerals of this au-cu series can contain high concentrations of pd and sometimes pt. the ag-enriched (au,cu) alloys were previously reported (nielsen et al. 2003a–e; cabri et al. 2005a; rudashevsky & rudashevsky 2005a, b, 2006a; rudashevsky et al. 2009a, b, 2010a–d, 2012a–i, 2014, 2015). grains of tetra-auricupride (aucu) are often hosted in cu-fe sulphide grains composed of chalcocite and bornite. a total of 1595 grains of au, cu, and ag compounds were found in 27 of the 30 samples using bse imagery and emp analysis. they are also observed in polished fig. 18 2 columns a b c d e f g h 24-1053, 40 #36 10 μm 18-978, 63 #23 10 μm18-978, 63 #28 10 μm 24 1057, 45-1 #179 10 μm 24-1045, 40 #103 10 μm bs0301 45 #139 30 μm bs0301, 45 #17 30 μm 18-978, 63 #5 10 μm fig. 18 bse images of free grains of cu and pd minerals in polished monolayer thin sections of hs concentrates. the rounded grains may be droplets or grains formed during exsolution in sulphide (e. makovicky, pers. comm. 2008). note the grain in panel f, which appears to have formed by amalgamation of two individual grains. abbreviated names of minerals, compounds and alloys are as follows: sk: skaergaardite (pdcu) including sk1: (pdcu)) and sk2: ((pd,au)cu). (cu,pd)α: unnamed alloy. ch: chalcocite. bn: bornite. table 3 provides formulae for the precious metal minerals. formulae for common rock-forming minerals and base metal sulphides are found in appendix 1. grain numbers and scales included in the images. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 29 of 66 www.geusbul let in.org thin sections of the host gabbro from au-rich mineralisation levels, such as in samples 24-1018 (rudashevsky & rudashevsky 2005a) and bs0304 (cabri 2003). the au-cu phases are identified based on their compositions. a total of 804 emp analyses identify six groups of cu-au phases on the basis of the relative proportion of (au+pd+pt+ag) (at.%; fig. 21). the groups are as follows: (1) 20–30 at.% auricupride (aucu3); (2) 40–45% (сu,au) alloys, which may be similar to cu-rich tetra auricupride of zaccarini et al. (2004); (3) 45–55 at.% tetra-auricupride (aucu); (4) 55–70 at.% (au,cu) alloys; (5) 70–80 at.% unnamed mineral (au3cu); (6) 80–100 at.% (au,cu,ag) alloys. native gold and silver are best referred to as (ag,au) and (ag,cu). grains of au, cu and ag were recovered from 27 samples and represent c. 30.2% of the total area of all precious metal grains found in the studied samples (table 3). tetra-auricupride (aucu) dominates with 19% (row 8), followed by (au,cu,ag,pd) alloys (row 13) with 5.8% and (au,cu,pd) alloys with 2.9% (row 11). all other au-rich compositions represent <0.75% and sum to a total of 2.3% (table 3, rows 9, 10, 12, 14–18). average compositions of all au, cu and ag minerals and compounds are listed in table 9 and the compositional range of au and fig. 19 2 columns a b c d e f g h i j k l bs0304 45 #185 10 μm bs0304 45 #290 10 μm 18-1010, 45-1 #77 10 μm 18-1010, 75-2 #1a 30 μm 18-1010, 75-2 #2 10 μm 24-1057, 45-1 #111 10 μm 18-958, 40 #3 10 μm 18-978, 63 #93 10 μm 24-1045, 40 #162 10 μm 24-1057, 45-1 #158 10 μm 18-1010, 45-1 #54 10 μm 24-1022, 40 #80 10 μm fig. 19 bse images showing the relationship between cu and pd minerals and other precious metal minerals in polished monolayer thin sections of hs concentrates. table 3 provides formulae for the precious metal minerals. formulae for common rock-forming minerals and base metal sulphides are found in appendix 1. note the high proportions of precious metal to cu-fe sulphide and in the rimming of skaergaardite by tetra-auricupride in panel i. abbreviated names of minerals, compounds and alloys are as follows: sk: skaergaardite. (cu,pd)n: alloy of cu and pd. hng: hongshiite (pt,pd,au) (cu,fe). vys: vysotskite. vsl: vasilite. at: atokite. kth: keithconnite. zv: zvyagintsevite. cbr: cabriite. (pd,cu,sn): alloys of pd, cu and sn. aucu and (au,pd) cu: tetra-auricupride. (au,pd)3cu: unnamed compound. ag: native silver. bn: bornite. ch: chalcocite. copn: cobaltian pentlandite. grain numbers and scales included in the images. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 30 of 66 www.geusbul let in.org fig. 20 bse images of pt minerals in polished monolayer thin sections of hs concentrates. table 3 provides formulae for the precious metal minerals. formulae for common rock-forming minerals and base metal sulphides are found in appendix 1. abbreviated names of minerals, compounds and alloys are as follows: (pt,cu,fe), (pt,fe,pd,cu), (pt,fe,pd), (pt,fe), (cu,pt): alloys of pt, cu, fe, and pd. (pt,cr): alloys of pt and cr. (au,cu): alloys of au and cu. sp: sperrylite. sk: skaergaardite. zv: zvyagintsevite. kth: keithconnite. apd: arsenopalladinite. vnc: vincentite. aucu: tetra-auricupride. au3cu: unnamed compound. bn: bornite. ch: chalcocite. cp: chalcopyrite. ilm: ilmenite. mt: magnetite. grain numbers and scales included in the images. fig. 20 1.5 columns 24-1057, 45-1 #168 10 μm 18-1010, 45-1 #114 10 μm a b c d e f g h i j k l 23a-807, 45-2 #39 10 μm 24-1022, 40 #97 10 μm 23a-808, 45-1-1 #64 10 μm bs0304 45 #278 10 μm 18-1012, 40 #77 10 μm 24-977, 40 #3 10 μm 24-1018, 40-2 #4 10 μm 23a-808, 45-1-1 #63 10 μm 23a-798, 40-1 #79 10 μm 23a-806, 45-1 #41 10 μm table 8 average composition of pt alloys row no. sample no. grain no. pt pd au fe cu ni cr sum (pt,fe,pd,cu) alloy 1 all mineralisation ave. of 18 grains wt% 70.5 12.9 0.4 11.7 3.4 0.1 n.d. 99.2 prop. 0.48 0.16 0.00 0.28 0.07 0.00 1.00 sd 9.46 6.70 1.17 1.20 1.91 0.20 (pt,cu,fe,pd) alloy 2 all mineralisation ave. of 21 grains wt% 52.3 17.5 2.2 11.2 15.8 0.2 n.d. 98.5 prop. 0.30 0.18 0.01 0.22 0.28 0.00 1.00 sd 7.57 8.77 1.93 0.91 1.83 0.29 (cu,pt) alloy 3 18-1012 40#77 wt% 28.6 1.0 n.d. 2.5 67.9 n.d. n.d. 100.0 prop. 0.12 0.01 0.04 0.83 1.00 (pt,cr,pd) alloy 4 bs0304 45#278 wt% 87.1 3.1 2.4 n.d. 0.6 0.7 4.6 98.5 prop. 0.75 0.05 0.02 0.02 0.02 0.14 prop.: atomic proportion. sd: standard variation. n.d.: not detected. indicates no data. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 31 of 66 www.geusbul let in.org cu minerals are presented in table 10. the alloys dominated by au and cu often contain significant concentrations of pd, occasionally pt and te (table 10). exotic (cu,au) alloys with wt% fe, zn and ni substitutions were also found (e.g. grain 24-1018 40#274 with au 68.5 wt%, pd 0.8 wt%, cu 24.3 wt%, fe 5.1 wt% and a total of 98.7 wt%, and grain 18-988 40#5 with au 59.8 wt%, ag 2.0 wt%, cu 24.2 wt%, fe 0.3 wt%, zn 5.8 wt%, ni 7.6 wt%, and a total of 99.8 wt%). the relative proportions of auand cu-alloys vary throughout the mineralised gabbros. in the central parts of mineralisation, the ordered compounds such as tetra-auricupride (aucu) and auricupride (aucu3) comprise 90.7% of all au-cu phases in all au-rich mineralisation levels for which data are available (pd1/au in drill cores 90-24 and 90-18 and bulk sample bs0304 from tof; fig. 22a). in sharp contrast, ag-rich alloys (au,cu,ag) dominate (up to 100%) in samples 90-23a-798 and 90-10-434, from the au-rich pd4 mineralisation level in drill cores nearer to the margins of the intrusion (fig. 22b). these paragenetic variations between pd-rich and au-rich mineralisation levels and between the centre and margin of the intrusion are paralleled in the compositions of the au-cu minerals (table 11). the average composition of tetra-auricupride (aucu) in pd5 is significantly enriched in pd (22.3 wt%) and pt (1.6 wt%; at 95% confidence level) compared to tetra-auricupride in the pd1/au layer (4.9 wt% pd, 0.4 wt% pt; table 11, rows 1 0 20 40 60 80 100 120 140 160 20–25 25–30 30–35 35–40 40–45 45–50 50–55 55–60 60–65 65–70 70–75 75–80 80–85 85–90 90–95 95–100 cu au (cu,au) cuau (au,cu) au+pd+pt+ag in cu+au+pd+pt+ag (at.%) (au,cu,ag,pd) (au,ag) 3 au n = 804 cu3 fig. 21 one column n um be r o f a na ly se s fig. 21 compositions and proportions of au-cu minerals and phases. the histogram shows the wide compositional range and the frequency of au-rich minerals, intermetallic compounds and alloys of au and cu in the mineralisation. based on compositions of a total of 804 analyses and expressed as sum of (au+pd+pt+ag) in at.%. the data set includes 0.1% unnamed compound au3cu and 0.1% (au,cu,ag) alloy, which are not shown here. abbreviations are as follows: cuau: tetra-auricupride. cu3au: auricupride. (au,cu), (cu,au), (au,cu,ag,pd): alloys of au, cu and ag. table 9 average compositions of au, cu, pd and ag minerals row no. no. of grains   au pd pt ag cu fe sum tetra-auricupride aucu 1 275 ave. (wt%) 63.9 8.7 0.6 n.d. 25.6 0.4 99.3 prop. 0.79 0.20 0.01 0.98 0.02 2.00 sd 10.60 8.20 1.37 2.11 0.49 auricupride aucu3 2 23 ave. (wt%) 46.8 4.1 n.d. n.d. 47.4 1.2 99.4 prop. 0.92 0.14 2.86 0.08 4.00 sd 5.91 3.97 2.05 0.79 unnamed au3cu 3 76 ave. (wt%) 76.3 11.4 1.2 n.d. 10.4 0.5 99.7 prop. 2.30 0.63 0.04 0.97 0.05 4.00 sd 9.76 8.12 2.30 1.52 0.55 (au,cu,pd) alloy 4 132 ave. (wt%) 78.9 6.24 0.12 n.d. 13.6 0.4 99.2 prop. 0.60 0.08 n.d. 0.31 0.01 1.00 sd 7.50 4.71 0.53 5.51 0.42 (cu,au,pd) alloy 5 41 ave. (wt%) 63.5 4.9 0.1 n.d. 30.1 0.5 99.1 prop. 0.38 0.05 0.00 0.56 0.01 1.00 sd 7.70 5.90 0.60 2.10 0.60 (au,cu,ag,pd) alloy 6 189 ave. (wt%) 89.5 1 0.2 2.7 5.6 0.6 99.6 prop. 0.77 0.02 0.00 0.04 0.15 0.02 1.00 sd 4.73 2.41 0.60 2.77 2.67 0.60 native gold, electrum (au,ag) 7 6 ave. (wt%) 76.5 n.d. n.d. 21.7 n.d. n.d. 99.2 prop. 0.66 0.34 1.00 sd 12.10 12.60 copper silver (ag,cu) 8 3 ave. (wt%) n.d. n.d. n.d. 93.4 5.8 0.3 99.5 prop. 0.90 0.10 1.00 sd 0.64 0.78 0.31 ave.: average. prop.: atomic proportions. sd: standard deviation. n.d.: not detected. indicates no data. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 32 of 66 www.geusbul let in.org and 2). alloys (au,cu,ag,pd) from pd5 in drill cores near the margins of the intrusion (samples 90-23a-806, -807, -808 and 90-10-443, -445) are enriched in pd, pt and ag, compared to the overlying au-rich pd4 mineralisation level (table 11, rows 3 and 4). the same type of compositional variation is shown in most of the au-cu minerals. they are pd-rich in pd5 and pd-poor in the au-rich mineralisation levels, irrespective of the absolute elevation of the au peak above pd5 (table 11, rows 5–12). 4.3.3.1 relationships of intermetallic compounds and (au,cu,ag) alloys to cu-fe sulphides the au-rich phases are mostly found at the margins of sulphide grains (fig. 23a). common droplet-like sulphide globules as well as irregular grains of sulphides host microdroplets of au-minerals. the unique textures of these types of ‘two-liquid’ grains include: 1. droplets of au-phases (au,cu) at the margins of droplet-like sulphide grains (fig. 23b, c, e) 2. metal droplets inside sulphide hosts (fig. 23d, k) 3. microdroplets of cu-sulphide inside larger droplets of precious metal (fig. 23e, f). in addition, au-rich minerals also form idiomorphic cubic and cuboctahedral crystals in sulphide hosts (e.g. au in fig. 23g and (cu,au) in fig. 23h). these occurrences of au phases are all related to sulphide grains (fig. 23) with few exceptions. near the margins of the intrusion in drill cores 90-10 and 90-23a, however, the au phases are commonly separated spatially from the sulphides. in these gabbros, the au phases are unattached to sulphides (fig. 23l–n) and commonly form elongated grains related to hydrous silicates in interstitial spaces (also illustrated in godel et al. (2014)). free grains of au minerals, as well as more irregular droplet-like particles of (au,cu,pd) and (au,cu,ag) table 10 compositional range of au-cu minerals a mineral type au pd pt ag cu fe te sum (wt%) tetra-auricupride aucu pd 44.4 23.3 1.9 n.d. 29.6 0.6 n.d. 99.8 pd-pt 36.9 24.0 8.5 n.d. 28.4 1.5 n.d. 99.3 ag 54.7 9.6 n.d. 8.8 25.8 0.8 n.d. 99.7 auricupride aucu3 pd 30.0 13.5 n.d. n.d. 51.1 3.5 n.d. 98.1 pt 22.6 4.9 22.3 n.d. 48.5 1.2 n.d. 99.5 unnamed au3cu pd 55.2 26.7 2.7 n.d. 13.6 1.1 n.d. 99.3 pt 51.7 17.9 18.6 n.d. 8.4 2.7 n.d. 99.3 pd-ag 80.8 6.7 n.d. 2.4 9.5 n.d. n.d. 99.4 (au,cu,pd) alloy te 49.1 23.2 2.7 n.d. 14.5 1.2 8.2 98.9 pt 61.3 14.7 3.6 n.d. 19.0 1.0 n.d. 99.6 (cu,au,pd,pt) alloy pd 59.8 16.1 n.d. n.d. 22.8 0.9 n.d. 99.6 pt 66.0 n.d. 3.5 n.d. 28.0 1.3 n.d. 98.8 (au,cu,ag ) alloy cu 79.0 2.3 n.d. 1.5 14.7 2.0 n.d. 99.5 ag 74.5 1.4 n.d. 17.9 3.2 1.1 n.d. 98.1 pd-pt 67.8 17.3 3.5 0.9 9.5 n.d. n.d. 99.0 pt 90.6 n.d. 3.6 2.0 3.8 n.d. n.d. 100.0 a emp analyses of individual grains with maximum substitutions n.d.: not detected. (cu,au) 3.4% aucu 89.3% cu₃au 1.4% (au,cu) 5.75% a (au,cu) 0.6% (au,cu,ag) 99.4% b fig. 22 one column fig. 22 au-rich parageneses in a: mineralisation level pd1/au in all drill core samples from the centre of the intrusion and b: mineralisation level pd4 near the margins of the intrusion. in (a), the average is based on data from samples 90-24-1018, 90-18-958 and bulk sample bs0304 from tof (locations in fig. 1) and is representative for the au-rich levels in the central parts of mineralisation. in (b) the average is based on samples 90-23a-798 and 90-10-434. aucu: tetra-auricupride; cu3au: auricupride; (au,cu), (cu,au), (au,cu,ag): alloys of au, cu and ag. not shown in (a): 0.1% unnamed compound au3cu and 0.1% (au,cu,ag) alloy. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 33 of 66 www.geusbul let in.org alloys and grains with compositions of tetra-auricupride (aucu), auricupride (cu3au) and the unnamed mineral au3cu (fig. 24a–j) are common in hs concentrates. more complex grains are referred to as micronuggets and may have formed by amalgamation of smaller grains. 4.3.3.2 grain intergrowths of more au-cu minerals au-rich grains from the central and upper au-rich pd2 and pd1 mineralisation levels are commonly zoned (fig. 25a–g; table 12). observed types of zoning include: 1. (au,cu) alloy rimmed by aucu (fig. 25a) 2. (cu,au) alloy rimmed by aucu3 followed by (au,ag,cu) alloy (fig. 25b) 3. (cu,au) alloy rimmed by (cu,ag,au) alloy (fig. 25c) 4. (cu,au) alloy rimmed by aucu3 (fig. 25d) 5. aucu rimmed by aucu3(fig. 25e) 6. aucu3 rimmed by aucu (fig. 25f, g) these types of zoned grains occur in hs concentrates in intergrowths with the sulphides bornite, chalcocite, chalcopyrite and as free grains. intergrowths of au and cu minerals with solid exsolution-like textures were also observed (fig. 25h, j; table 13) and include (au,cu,pd) alloy with lamellae of tetra-auricupride (aucu), (cu,au) alloy with auricupride (aucu3), and unnamed au3cu mineral with tetra-auricupride (aucu). 4.3.3.3 relationships of intermetallic compounds and (au,cu,ag) alloys with pges intergrowths between tetra-auricupride (aucu) and tellurides are very common in the au-rich pd1/au mineralisation level (fig. 26a, b). the most common telluride is kotulskite (pdte), followed by merenskyite (pdte2), melonite (nite2, fig. 26e) and keithconnite (pd3-xte; fig.  26c). small inclusions of arsenides can also be found in tetra-auricupride and include the minerals sperrylite (ptas2; fig. 26f) and orcelite (ni5-xas2; fig. 25d, j). in pd5 and pd4 mineralisation levels from drill cores 90-10 and 90-23a from near the margins of the intrusion, (au,cu,ag,pd) alloys are associated with pd arsenides and pt arsenides such as arsenopalladinite ((pd,cu)8as3; fig. 26g, h), palladoarsenide (pd2as; fig. 20h, j), and sperrylite (ptas2; fig. 20i, k); vincentite (pd,cu)3as (fig. 26h–j) and isomertieite pd11as2(sb,sn)2 (fig. 32h, i). in the au-rich pd4 horizon, kotulskite (pdte) often occurs in intergrowths with (au,cu,ag,pd) alloy (fig. 26h). cu-au minerals are rare in the skaergaardite dominated pd6 and pd5 mineralisation levels in central drill cores and sampling sites. here, the cu-au phases are associated with skaergaardite, keithconnite (pd3-xte; fig. 26k, l), and the pd sulphides vasilite ((pd,cu)16s7) and vysotskite ((pd,cu,ni)s; fig. 26m). in contrast, cu-au minerals comprise up to 28% of the parageneses in the uppermost metre of pd5 in drill cores near the margins of the intrusion (nielsen et al. 2003e, rudashevsky et  al. 2012d). here, the cu-au phases are related to pd-intermetallic compounds, arsenides and sulphides such as zvyagintsevite (pd3pb), atokite (pd3sn), keithconnite (pd3-xte), vysotskite ((pd,ni)s), arsenopalladinite (pd8as3), vincentite ((pd,cu)3as), palladoarsenide (pd2as), and sperrylite (ptas2). table 11 average compositions of au and cu minerals per mineralisation level row no. mineralisation level no. of grains au pd pt ag cu fe sum (wt%) tetra-auricupride aucu 1 pd5 38 46.8 22.3 1.6 n.d. 27.5 0.8 99.0 2 pd1/au 159 68.6 4.9 0.4 n.d. 25.2 0.3 99.4 (au,cu,ag,pd) alloy 3 pd5 30 84.6 4.9 0.7 4.6 3.4 0.6 98.8 4 pd4 159 90.5 0.2 0.1 2.3 6.1 0.5 99.6 auricupride aucu3 5 pd2 4 43.2 7.6 n.d. n.d. 47.7 1.0 99.5 6 pd1/au 20 47.7 3.0 n.d. n.d. 47.5 1.2 99.4 unnamed au3cu 7 pd5 12 64.4 21.5 1.1 n.d. 11.9 0.4 99.3 8 pd1/au 25 85.9 3.2 n.d. n.d. 10.1 0.1 99.3 (au,cu,pd) alloys 9 pd5-pd2 58 74.4 7.9 n.d. n.d. 16.8 0.4 99.5 10 pd1/au 11 84.0 1.3 n.d. n.d. 13.0 0.7 99.0 (cu,au,pd) alloys 11 pd2 7 61.1 11.9 n.d. n.d. 26.1 0.3 99.3 12 pd1/au 10 67.5 1.4 n.d. n.d. 28.8 1.3 99.0 n.d.: not detected. statistical means at the 95% confidence levels. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 34 of 66 www.geusbul let in.org 4.3.3.4 relationships between au-cu compounds and rock-forming minerals intermetallic compounds and alloys of au and cu occur in intergrowths with the primary rock-forming minerals pyroxenes, plagioclase, feti oxides and olivine (fig. 27a–j). near the margins of the intrusion in drill cores 90-10 and 90-23a, au-cu minerals are also found in veins between grains of cu-fe sulphides and typical lateto post-magmatic hydrous minerals, such as hornblende, actinolite, biotite, chlorite and talc (fig. 27k–p). 4.3.4 pge sulphides and s-bearing ag minerals pge sulphides account for 3.8% of all studied pges and au-rich grains (table 3, rows 23–33). they are found at all mineralisation levels but are most common in drill cores from the western part of the intrusion, such as in samples 90-18-988, 90-18-972, 90-10-445 and 90-10443, where pd, cu and ag sulphides constitute 40–70% of the precious metal parageneses (see nielsen et al. 2019a). their relative proportion is reduced to <1% in the au-rich mineralisation levels of cores 90-10 and 90-18 (fig. 10; see also nielsen et al. 2019a, table 4 and fig. 9). vysotskite ((pd,ni,cu)s) and vasilite ((pd,cu)16s7) are the most common pge-sulphides and account for 1.7% and 1.6% of all precious metal phases, respectively (fig. 28). the compositional ranges and the average compositions of a total of 100 grains of vysotskite and of 52 grains of vasilite are shown in table 14. in some cases, the vysotskite grains divide into two contrasting subgrains, with 3.6 wt% and 22.5 wt% pt, respectively (fig. 28d; table 14, rows 2 and 3). the remaining 0.5% include exotic sulphides rich in ag (ag,pd,cu), cu (cu,pd,ag) or fig. 23 1.5 columns a b c d e f g h i j k l 24-1018, 63 #17 30 μm m n 24-1018 ps 30 μm bs0304 75 #4 30 μm bs0304 75 #52 30 μm bs0304 75 #18 30 μm 24-1022, 40 #83 10 μm 24-1018, 63-2 #5 10 μm bs0304 75 #36 30 μm 24-1022, 63 #19 10 μm 23a-807, 125 #2 30 μm 10-434, 40-2 #7 10 μm 10-434, 40-2 #10 10 μm 23a-798, 40-1 #24 10 μm 23a-798, 40-1 #144 10 μm fig. 23 bse images showing the petrographic relations between (au,cu) phases and rock-forming minerals of the gabbroic host and sulphides. a: polished thin section of gabbro with interstitial sulphide and tetra-auricupride (aucu). b–n: polished monolayer thin sections of hs concentrates. panels a–i are from samples in the central part of the intrusion, and j–n are from samples near the margins of the intrusion. table 3 provides formulae for the precious metal minerals. formulae for common rock-forming minerals and base metal sulphides are found in appendix 1. grain numbers and scales included in the images. abbreviated names of minerals, compounds and alloys are as follows: aucu: tetra-auricupride. (cu,au), (au,cu,ag): alloys of au, cu and ag. (au,pd)3cu: unnamed au3cu. kt: kotulskite (pdte). bn: bornite. ch: chalcocite. cp and cp: chalcopyrite. opx: orthopyroxene. cpx: clinopyroxene. pl: plagioclase. ilm: ilmenite. hb: hornblende. act: actinolite. kth: keithconnite. mt: magnetite. chl: chlorite. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 35 of 66 www.geusbul let in.org fig. 24 bse images of free grains of au and cu minerals in monolayer polished thin sections of hs concentrates. abbreviated names of minerals, compounds and alloys are as follows: aucu: tetra-auricupride. aucu3: auricupride. au3cu: unnamed alloy. (au,cu,pd), (au,cu,ag): alloys of au, cu, pd and ag. sp: sperrylite. kt: kotulskite. table 3 provides formulae for the precious metal minerals. formulae for common rock-forming minerals and base metal sulphides are found in appendix 1. grain numbers and scales included in the images. fig. 24 2 columns a b c d e f g h i j k 24-1034, 40 #61 10 μm 24-1034, 40 #156 10 μm 23a-798, 40-1 #97 10 μm 10-434, 40-2 #5 10 μm 24-1034 ,40 #74 10 μm 24-1018, 40 #221 10 μm 24-1018, 40 #37 30 μm 24-1018, 40 #138 10 μm bs0304 45 #258 10 μm 24-1018, 40 #78 10 μm 23a-798, 40-1 #92, 93 10 μm table 12 compositional zonation in au-cu mineral grains sample no. grain no. type of zoning phase au pd ag cu fe sum (wt%) multiple ave. of 10 grains (au,cu) → aucu (au,cu) 87.1 1.8 n.d. 10.8 0.2 99.9 aucu 73.2 1.8 n.d. 23.3 0.5 99.3 multiple ave. of 11 grains (cu,au) → aucu3 (cu,au) 64.6 4.2 n.d. 30.0 0.3 99.1 aucu3 47.8 3.0 n.d. 47.4 1.2 99.4 bs0304 45#243 (cu,au) → aucu3 → (au,ag,cu) (cu,au) 6.03 4.2 n.d. 31.8 0.5 99.5 aucu3 49.5 3.4 n.d. 44.8 1.5 99.2 (cu,ag,au) 54.5 2.3 26.3 13.9 2.0 99 bs0304 125#12 (cu,au) → (cu,ag,au) (cu,au) 67.4 2.2 n.d. 29.2 n.d. 98.8 (cu,ag,au) 20.1 n.d. 41.8 36.7 0.9 99.5 ave.: average. n.d.: not detected. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 36 of 66 www.geusbul let in.org table 13 composition of au-cu minerals and their exsolution phases sample no. grain no. phase au pd cu fe sum (wt%) 24-1034 ave. of 13 pairs host lamellae (au,cu,pd) 78.5 6.0 14.3 0.3 99.1 aucu 70.6 6.9 21.5 0.3 99.3 24-1018 40#95 host lamellae (cu,au) 65.0 1.6 31.6 1.3 99.5 aucu3 51.7 n.d. 45.4 2.3 99.4 24-1043 40#70 host lamellae au3cu 68.6 19.5 11.5 n.d. 99.6 aucu 54.9 18.4 25.4 0.4 99.1 ave.: average. n.d.: not detected. fig. 25 2 columns a b c d e f g h i j bs0304 75 #1 30 μm bs0304 45 #243 10 μm bs0304 125 #12 30 μm 24-1018, 63 #24 10 μm bs0304 45 #245 10 μm bs0303 45 #327 10 μm 24-1034, 40 #138 10 μm 24-1034, 40 #135 10 μm 24-1034, 40 #70 10 μm 24-1018, 40 #223 10 μm fig. 25 bse images of zoned grains of au minerals, cu minerals and exsolution textures in polished thin sections of hs concentrates. note the solid state exsolution of auricupride (panel j). table 3 provides formulae for the precious metal minerals. formulae for common rock-forming minerals and base metal sulphides are found in appendix 1. grain numbers and scales included in the images. abbreviated names of minerals, compounds and alloys are as follows: au3cu: unnamed compound. aucu: tetra-auricupride. aucu3 and cu3au: auricupride. (au,cu), (cu,au), (au,cu,pd), (cu,ag,au), (au,ag,cu): alloys of au, cu, ag and pd. zv: zvyagintsevite. kth: keithconnite. orc: orcelite (ni5-xas2). bn: bornite. and ch: chalcocite. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 37 of 66 www.geusbul let in.org (au,cu,pd) fig. 26 2 columns a b c d e f g h i j k l m n o bs0304 45 #321 10 μm 24-1018, 40 #39 10 μm 18-958, 40 #16 10 μm 24-1034, 40 #128 10 μm 24-1018, 40 #77 10 μm 24-1018, 40 #247 10 μm 23a-798, 40-1 #3 10 μm 10-434, 100 #1 10 μm 10-434, 40 #16 10 μm 10-434, 40-2 #1 10 μm 24-1057, 45-1 #117 10 μm 24-1022, 40 #51 10 μm 18-1010, 45-1 #92 10 μm 23a-808, 45-1-1 #61 10 μm 23a-807, 45-2 #67 10 μm fig. 26 bse images of the petrographic relations between au minerals and cu minerals with pges in polished monolayer thin sections of the heavy mineral hs concentrates. the au and cu minerals are associated with virtually all other precious metal phases of the mineralisation. table 3 provides formulae for the precious metal minerals. formulae for common rock-forming minerals and base metal sulphides are found in appendix 1. grain numbers and scales included in the images. abbreviated names of minerals, compounds and alloys are as follows: aucu: tetra-auricupride. (au,pd) cu: tetra-auricupride. (au,pd)3cu: unnamed compound. (au,cu,pd), (au,cu,ag): alloys of au, cu, pd and ag. (au,ag): native gold. sk: skaergaardite (pdcu). kt: kotulskite (pdte). kth and kth: keithconnite (pd3-xte). mln: melonite (nite2). sp: sperrylite. apd: arsenopalladinite. plas: palladoarsenide. vnc: vincentite. vys: vysotskite. vsl: vasilite. zv: zvyagintsevite. bn: bornite. ch: chalcocite. cp: chalcopyrite. chl: chlorite. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 38 of 66 www.geusbul let in.org fig. 27 2 columns a b c d e f g h i j k l m n o p 24-1022, 40 #84 30 μm 10-434, 40 #28 10 μm 24-1018, 40-2 #3 30 μm 18-958, 40 #4 10 μm 10-434, 70 #3 10 μm 23a-798, 40-1 #2 10 μm bs0304 45 #242 10 μm 10-434 ,40 #3 10 μm 24-1018, 40 #255 30 μm 24-1018, 63 #38 30 μm 23a-806, 75-1 #9 10 μm 23a-807, 45-1 #6 10 μm 23a-798, 40-2 #4 10 μm 23a-806, 45-1 #24 10 μm 23a-798, 63-1 #2 10 μm 23a-806, 45-2 #5 10 μm fig. 27 relationship between au and cu minerals and rock-forming minerals of the gabbroic host in polished monolayer thin sections of hs concentrates. note that au-cu phases in general coexist with ‘dry’ parageneses in the centre of the intrusion (samples from drill cores 90-24 and 90-18, and bulk sample bs0304), and with hydrous silicate parageneses near the margins (samples from drill core 90-10 and 90-23a). table 3 provides formulae for the precious metal minerals. formulae for common rock-forming minerals and base metal sulphides are found in appendix 1. grain numbers and scales included in the images. abbreviated names of minerals, compounds and alloys are as follows: aucu: tetra-auricupride. (au,cu), (au,cu,ag), (au,pd,cu), (au,cu,pd): alloys of au, cu, pd and ag. au3cu and (au,pd)3cu: unnamed compounds. apd: arsenopalladinite (pd8as3). zv: zvyagintsevite (pd3pb). kth: keithconnite (pd3-xte). kt: kotulskite (pdte). cpx: clinopyroxene. opx: orthopyroxene. pyr: pyroxene. pl: plagioclase. ol: olivine. ilm: ilmenite. hb: hornblende. act: actinolite. bt: biotite. chl: chlorite. tlc: talc. bn: bornite. ch: chalcocite. dgn: digenite. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 39 of 66 www.geusbul let in.org pd (pd,ag,cd,hg,tl) (table 3, rows 25–32; fig. 29). the substituting elements are the formulae listed in order of decreasing proportion. single grains of stephanite (ag5sbs4), polybasite (ag16sb2s11) and acanthite/argentite (ag2s; fig. 29a–c) were also found and are all near-stoichiometric. compositions of very rare and exotic unnamed sulphides composed of pd, cu, pt, ag, cd, hg and tl (table 15) include coldwellite (pd3ag2s; mcdonald et al. 2015; fig.  29e; table 15, row 1). irregular and droplet-like inclusions of the sulphides include ideal droplets of bornite, chalcocite and vysotskite (pds) in ilmenite (fig. 28a), a droplet of bornite and vysotskite in a bornite host (fig. 28b) and microdroplets of vysotskite and vasilite (pd16s7) in bornite and chalcopyrite-bornite host globules (fig. 28c–i). polymineralic grains of vysotskite and vasilite with other pd-intermetallic compounds occur in samples near the margins of the intrusion. the intermetallic compounds include zvyagintsevite (pd3pb), keithconnite (pd3-xte) and arsenides, such as arsenopalladinite (pd8as3), palladoarsenide (pd3as) and vincentite ((pd,pt)3(as,sb,te)), in association with cu-fe sulphides and hydrous silicates (e.g. fig. 26o). fig. 28 2 columns a b c d e f g h i j k 18-1010, 75-2 #29 10 μm 18-972, 125 #6 30 μm 24-1056, 125 #3 30 μm 18-988, 40 #25 10 μm 18-1010, 45-1 #41 10 μm 10-443, 40 #35 10 μm 10-445, 40 #1 10 μm 18-1001, 40 #8 10 μm 18-1010, 45-1 #11a 10 μm 18-988, 40 #24 10 μm 18-1010, 45-2 #4 10 μm fig. 28 bse images of vysotskite and vasilite in polished monolayer thin sections of hs concentrates. the precious metal sulphides form as exsolutions in sulphide droplets, as exsolutions on the margins of sulphide globules and as irregular grains intergrown with other precious metal phases. table 3 provides formulae for the precious metal minerals. formulae for common rock-forming minerals and base metal sulphides are found in appendix 1. grain numbers and scales included in the images. abbreviations: vys: vysotskite, including pd-rich vys1 and pt-rich vys2. vys-(pt): vysotskite with no pt. vsl: vasilite. kth: keithconnite. (au,pd)cu: tetra-auricupride. au3cu: unnamed compound. bn: bornite. ch: chalcocite. cp: chalcopyrite. dgn: digenite. pn: pentlandite. ilm: ilmenite. chl: chlorite. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 40 of 66 www.geusbul let in.org table 14 compositional types of vysotskite and vasilite in the skaergaard pge-au mineralisation row no. sample no. grain no. type pd pt cu fe ni s te sum vysotskite (pd,ni,cu,pt)(s,te) 1 18-1010 45#106 ‘pure’ a wt% 75.0 n.d. n.d. n.d. 0.9 23.0 n.d. 98.9 prop. 0.98 0.02 1.00 2.00 2 18-988 40#18 ni-pt wt% 66.8 3.6 1.0 n.d. 5.0 23.9 n.d. 100.2 prop. 0.84 0.02 0.04 0.11 1.00 2.00 3 18-988 40#18 pt wt% 48.2 22.5 1.1 0.9 2.5 21.1 n.d. 99.4 prop. 0.69 0.20 0.03 0.03 0.06 1.00 2.00 4 18-988 40#18 cu wt% 65.3 n.d. 8.6 1.2 n.d. 24.2 n.d. 99.3 prop. 0.8 0.18 0.03 n.d. 0.99 2.00 5 18-1010 45-46 ni wt% 63.1 n.d. 1.3 0.7 8.6 25.2 n.d. 98.9 prop. 0.76 0.03 0.02 0.19 1.01 2.00 6 18-1010 75/2#29 te wt% 33.6 33.7 6.1 1.6 n.d. 17.2 6.4 98.6 prop. 0.52 0.29 0.16 0.05 n.d. 0.9 0.08 2.00 7 multiple ave. of 100 grains wt% 71.1 0.8 1.5 0.6 1.8 23.6 0.1 99.5 prop. 0.91 0.01 0.03 0.01 0.04 1.00 2.00 sd 5.75 4.31 1.48 0.45 2.41 0.95 0.64 vasilite (pd,cu,fe)16s7 8 multiple ave. of 52 grains   wt% 72.2 14.0 0.5 n.d. n.d. 12.7 n.d. 99.4 prop. 11.97 3.88 0.17 6.99 23.00 sd 1.26 0.96 0.41 0.20 a close to empirical formula pds. ave.: average. prop: atomic proportions. sd: standard deviation. n.d.: not detected. indicates no data. fig. 29 2 columns a b c d e f g h 18-1012, 40 #99 10 μm 18-988, 40 #4 10 μm 18-972, 125 #4 30 μm 24-1018, 40 #36 30 μm 23a-806, 45-1 #45 30 μm 23a-807, 45-2 #25 30 μm 24-1018, 40 #111 10 μm 24-1022, 40 #93 10 μm fig. 29 bse images of rare pge sulphides in polished monolayer thin sections of hs concentrates. all phases are rare and some are found only as single grains. table 3 provides formulae for the precious metal minerals. formulae for common rock-forming minerals and base metal sulphides are found in appendix 1. formulae are given here for rare minerals that are not listed in table 3 or appendix 1. abbreviated names of minerals, compounds and alloys are as follows: aucu: tetra-auricupride. ag2s: acanthite/argentite. stph: stephanite (ag5sbs4). plbs: polybasite (ag16sb2s11). pd3ag2s: coldwellite. sk: skaergaardite. un3: unnamed compounds, including (pd3cu6s3te2), (pd,hg,ag)2s, (cu,pd)2-xs, (pd,cu,pt)3s2 and (pd,ag,cd,cu,tl)4s. zv: zvyagintsevite. vys: vysotskite. bn: bornite. ch: chalcocite. dgn: digenite. act: acanthite. chl: chlorite. grain numbers and scales included in the images. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 41 of 66 www.geusbul let in.org these pd sulphides appear to form one single paragenesis with cu-pd intermetallic compounds and alloys, mostly with skaergaardite (cupd; fig. 19b–d), and with au and cu minerals (figs 26m–o, 28e, j) in cu-fe sulphide globules together with keithconnite, zyagintsevite and other pges. 4.3.5 intermetallic compounds of pd, cu and ag with pb, te, sn, sb and bi intermetallic compounds of pd, cu and ag with pb, te, sn, sb and bismuth (bi) are widespread in the mineralisation (table 3, rows 40–54), but account for no more than 2.6% of all observed precious metal phases. they include 1.6% zvyagintsevite (pd3pb), 0.5% keithconnite (pd3-хte), 0.2% atokite (pd3sn), 0.1% kotulskite (pd(te,bi)2-x; fig.  30a–i) and 0.18% non-stoichiometric (pd,cu,sn,pb,te) alloys (table 3, row 52). the group includes some very rare phases, sometimes represented by only a single grain, such as cabriite (pd2cusn), merenskyite (pdte2), sopcheite (ag4pd3te4), telargpalite ((pd,ag)3(te,pb)), hessite (ag2te), naldrettite (pd2(sb,sn)), froodite (pdbi2), unnamed (pd2te), (pd2cutebi) and ((pd,ag)2te). they represent 0.5% of the total precious metal paragenesis in this study (fig. 31). intermetallic compounds of pd, cu, and ag with pb, te and sn are found in all mineralisation levels. they are common in the samples from the marginal part of the intrusion, for example in drill core 90-23a, samples 23a-808 and 807 where they reach 69% and 60% respectively (nielsen et al. 2003b, e) . they reach 35% in sample 10-443 from the sulphide dominated drill core 90-10 (rudashevsky et al. 2012d). they are rare in the centre of the mineralisation (<1.5%; fig. 10). the te phases show a stratigraphic variation from keithconnite, typical for pd5, to te-rich kotulskite, merenskyite and non-stoichiometric (pd,cu,sn,pb,te) alloys in the au-rich pd1/au levels of central drill cores. zvyagintsevite (pd3pb) dominates in samples from mineralisation level pd5 of drill core 90-23a from near the eastern margin at the palaeogene lavas. in the lower part of pd5, zvyagintsevite accounts for 73% of the precious metal paragenesis and forms crystals up to 69 μm in size (nielsen et al. 2003e). it is closely associated with individual grains of keithconnite (pd3-xte) and atokite (pd3sn), with pd sulphides and arsenides, as well as with (au,cu,ag,pd) alloys in aggregates with cu-fe sulphides and h2o-bearing minerals including biotite, actinolite, chlorite and talc (figs 26o, 27p, 30b–d, h). in pd5 samples from central parts of the mineralisation, zvyagintsevite, occurs as rare and small grains <10 μm. they nucleated on skaergaardite (pdcu) globules in association with vasilite (pd16s7) and vysotskite (pds; fig. 19d). like all other precious metal minerals, these intermetallic compounds show compositional variations that reflect the local geochemical environment. the variation in compositions and the average for a total of 145 grains of zvyagintsevite from all studied samples are found in table 16, for 63 analyses of keithconnite and 13 analyses of kotulskite in table 17, for 14 analyses of atokite and 5 of cabriite (pd2cusn) in table 18 and rare intermetallic compounds in table 19. atokite (pd3sn) is associated to zvyagintsevite (figs 19g, 30h–j), but is rare, especially in drill cores (9023a, 0.2% of the paragenesis; table 3). keithconnite (pd3-xte) is found in 25 out of the 30 samples of this study, but only in very small proportions (0.5% of all pge and ag-au minerals; table 3). keithconnite was not found in the samples from the au-rich pd1/au mineralisation level from the centre of the intrusion. table 15 compositions of rare precious metal sulphides row no. sample no. grain no.   pd au pt ag cu fe te sn hg tl cd s sum coldwellite pd3ag2s 1 18-988 40#22 wt% 55.1 n.d. n.d. 37.5 1.0 n.d. n.d. n.d. n.d. n.d. n.d. 5.7 99.3 prop. 2.93 1.97 0.09 1.00 6.00 unnamed cu5fepd3(te,sn)2s3 2 24-1018 40#111 wt% 31.1 1.4 n.d. n.d. 32.0 5.3 16.2 5.8 n.d. n.d. n.d. 9.2 101.0 prop. 3.01 0.07 5.19 0.97 1.31 0.50 2.94 14.00 unnamed (cu,pd)2+xs 3 18-1012 40#99 wt% 11.5 n.d. 2.2 n.d. 68.2 n.d. n.d. n.d. n.d. n.d. n.d. 16.2 99.6 prop. 0.19 0.02 1.87 0.88 3.00 unnamed (cu,fe)(pd,pt)2s2 4 18-1012 40#99 wt% 37.4 n.d. 31.2 n.d. 10.6 n.d. n.d. n.d. n.d. n.d. n.d. 15.8 99.6 prop. 1.43 0.65 0.68 2.00 5.00 unnamed (pd,cu,fe)3(ag,cd,tl)s 5 18-988 40#4 wt% 58.8 n.d. n.d. 15.4 3.2 n.d. n.d. n.d. n.d. 6.9 7.8 6.4 99.3 prop. 2.77 0.39 0.25 0.17 0.35 1.00 5.00 unnamed (pd,hg,ag)2s 6 18-972  125#4 wt% 59.9 n.d. n.d. 7.3 1.6 n.d. n.d. n.d. 17.1 n.d. n.d. 12.1 98.6 prop. 1.49 0.18 0.07 0.23 1.00 3.00 prop.: atomic proportions. n.d.: not detected. indicates no data. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 42 of 66 www.geusbul let in.org keithconnite (pd3-xte) forms irregular grains on skaergaardite usually <10 μm, but up to 25 μm in size (fig. 19b– d, k). it is closely associated with au minerals, such as tetra-auricupride, unnamed (au3cu) and (au,cu,pd,ag) alloys, zvyagintsevite atokite, vysotskite, vasilite and pd-arsenides (figs 19j, k, 25h, 26c, d, o, 27l, p, 30e–g). keithconnite is also associated with cu-fe sulphides, and is observed in drill cores near the margins, commonly also associated to or intergrown with hydrous minerals, such as actinolite, talc and chlorite (figs 27l, p, 30e, f). 4.3.6 arsenides of pd, pt and ni arsenides of pd, pt and ni represent c. 2.1% of all precious metal minerals of the mineralisation (table 3, rows 33–38). the arsenide grains have irregular outlines and are 10–20 μm in size. the largest grain has an ecd of 69 μm (sperrylite in table 3). in general, they are associated with the pd minerals zvyagintsevite, keithconnite, atokite and kotulskite, the pd sulphides vysotskite and vasilite, the au minerals (au,cu,ag,pd) alloys, unnamed au3cu, tetra-auricupride, rare native gold, cu-fe sulphides and hydrous silicates, such as biotite, hornblende, actinolite and talc (figs 20h–l, 24j, 25d, j, 26f–j, n, 27f, 30d, k, l, 32). in some cases, vincentite forms rims to zvyagintsevite (fig. 30d). arsenopalladinite (pd8as3) represents 1%, followed by vincentite ((pd,pt)3(as,sb,te)) with 0.9% (table 3). the identification of arsenopalladinite is based on optical fig. 30 2 columns a b c d e f g h i j k l 23a-807, 75 #7 30 μm 23a-808, 45-1-2 #14 10 μm 23a-807, 45-2 #23 10 μm 24-1045, 40 #22 10 μm 23a-807, 45-1 #5 10 μm 23a-806, 45-1 #4 30 μm 23a-807, 45-2 #14 10 μm 23a-807, 45-1 #23 10 μm 18-1010, 45-1 #3 10 μm 18-978, 63 #44 10 μm 10-443, 40 #1 10 μm 23a-798, 80 #3 10 μm fig. 30 bse images of pd, cu and ag intermetallic compounds with pb, te, sn, sb and bi in polished monolayer thin sections of hs concentrates. zvyagintsevite (pd3pb) is the dominant pd phase in samples from the eastern margin, commonly in association with actinolite. table 3 provides formulae for the precious metal minerals. formulae for common rock-forming minerals and base metal sulphides are found in appendix 1. abbreviated names of minerals, compounds and alloys are as follows: zv: zvyagintsevite. kth: keithconnite. kt: kotulskite. vnc: vincentite. gng: guanglinite (obsolete and renamed to vincentite). apd: arsenopalladinite. at: atokite. sk: skaergaardite. vys: vysotskite. vsl: vasilite. (au,pd,cu) and (au,cu,ag): alloys au, cu and pd. bn: bornite. cp: chalcopyrite. ch: chalcocite. ilm: ilmenite. mt: magnetite. act: actinolite. chl: chlorite. grain numbers and scales included in the images. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 43 of 66 www.geusbul let in.org fig. 31 1.5 columns a b c d e f g h i 24-1034, 40 #161 10 μm bs0304 45 #203 10 μm 24-1034, 40 #115 10 μm 24-1022, 40 #81 30 μm 10-434, 40 #5 10 μm 24-1024, 80 #2 10 μm 24-1048, 40 #24 10 μm 24-1022, 63 #18 10 μm 18-978, 40 #29 30 μm fig. 31 bse images of rare intermetallic compounds of precious metals in polished monolayer thin sections of hs concentrates. the rare occurrence of these phases illustrates parageneses buffered by very localised geochemical conditions. table 3 provides formulae for the precious metal minerals. formulae for common rock-forming minerals and base metal sulphides are found in appendix 1. formulae are given here for rare minerals that are not listed in table 3 or appendix 1. abbreviated names of minerals, compounds and alloys are as follows: cbr: cabriite (pd2cusn). mrn: merenskyite (pdte2). nld: naldrettite (pd2(sb,sn)). tlp: telargpalite ((pd, ag)3te). hs: hessite (ag2te). sop: sopcheite (ag4pd3te4). fr: froodite (pdbi2). pd2te, (pd,ag)2te and pd2cutebi: unnamed compounds. sk: skaergaardite. aucu: tetra auricupride. (cu,pd)α: unnamed compound. (au,ag): native gold. au3cu: unnamed compond. (pt,fe,cu): alloy of pt, fe and cu. bn: bornite. cp: chalcopyrite. copn: cobaltian pentlandite. ch: chalcocite. opx: orthopyroxene. grain numbers and scales included in the images. table 16 compositional types of zvyagintsevite in the skaergaard au-pge mineralisation sample no. grain no. type   pd pt au cu fe sn te pb as sum 23а-808 45/1#11 ‘pure’ a wt% 60.5 n.d. n.d. n.d. n.d. n.d. n.d. 39.1 n.d. 99.6 prop. 3.00 1.00 4.00 23a-808 45/1#41 pt wt% 56.6 6.8 n.d. n.d. 0.7 2.8 n.d. 31.9 n.d. 98.8 prop. 2.81 0.18 0.07 0.12 0.81 4.00 23a-807 45/2#67 au wt% 52.6 n.d. 9.4 0.9 n.d. n.d. n.d. 33.8 1.3 98.0 prop. 2.68 0.26 0.08 0.89 0.09 4.00 bs0301 75/2#16 cu wt% 55.6 n.d. n.d. 3.6 0.6 n.d. n.d. 40.7 n.d. 100.5 prop. 2.66 0.28 0.06 1.00 4.00 23a-806 45/1#28 zone 1 wt% 57.0 2.5 1.6 n.d. 0.7 2.2 n.d. 35.5 n.d. 99.5 prop. 2.82 0.07 0.04 0.07 0.10 0.90 4.00 23a-806 45/1#28 zone 2 wt% 55.4 n.d. n.d. n.d. 9.2 n.d. 9.9 20.8 4.6 99.9 prop. 2.24 0.71 0.34 0.44 0.27 4.00 23a-808 45-1#28 sn wt% 62.4 1.2 3.2 n.d. n.d. 10.4 n.d. 22.0 n.d. 99.2 prop. 2.91 0.03 0.08 0.44 0.53 4.00 23a-808 45-1#15 as wt% 66.4 n.d. n.d. n.d. 1.2 n.d. n.d. 23.6 7.5 98.7 prop. 2.90 0.10 0.53 0.47 4.00 multiple   ave. of 145 grains   wt% 58.2 0.8 1.0 0.2 0.5 0.4 0.2 37.3 0.2 99.4 prop. 2.90 0.02 0.03 0.02 0.04 0.02 0.01 0.95 0.02 4.00 sd 2.51 1.33 1.68 0.51 0.85 1.50 1.29 5.10 1.04 a close to empirical formula pd3pb. ave.: average. prop.: atomic proportion. sd: standard deviation. n.d.: not detected. indicates no data. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 44 of 66 www.geusbul let in.org properties and emp analysis only. rare sperrylite (ptas2) accounts for 0.07% and palladoarsenide (pd2as) 0.06% of all precious metal phases (table 3). a few grains of isomertieite (pd11as2(sb,sn)2), majakite (pdnias) and unnamed pd11as2(sn,sb)2 were also recorded. representative compositions of the arsenides and their average compositions are listed in tables 20 and 21 and include data from 67 grains of arsenopalladinite, 6 grains of palladoarsenide, 64 grains of vincentite, 5 grains of isomertieite, a single grain of majakite and one of unnamed pd11as2(sn,sb)2. table 17 compositional types of keithconnite and kotulskite sample no. grain no. type   pd pt au cu fe sn te pb as sb bi sum keithconnite pd3-x(te,sn,pb) bs0304 45#240 ‘pure’ a wt% 63.0 n.d. n.d. 1.2 0.6 n.d. 32.4 1.1 n.d. n.d. n.d. 98.3 prop. 2.69 0.09 0.05 1.15 0.02 4.00 23a-806 45/2#7 au wt% 62.8 n.d. 2.6 n.d. 0.7 n.d. 28.8 4.9 n.d. n.d. n.d. 99.8 prop. 2.73 0.06 0.06 1.04 0.11 4.00 18-1012 125#1 cu wt% 63.6 n.d. n.d. 5.0 n.d. n.d. 30.1 1.4 n.d. n.d. n.d. 100.1 prop. 2.60 0.34 1.03 0.03 4.00 18-958 40#5 pb wt% 60.6 n.d. n.d. 1.1 1.1 n.d. 29.2 6.4 n.d. n.d. n.d. 98.4 prop. 2.63 0.08 0.09 1.06 0.14 4.00 18-972 40#30 sn wt% 70.2 n.d. n.d. n.d. 1.6 7.0 16.6 n.d. 3.8 n.d. n.d. 99.3 prop. 2.84 0.13 0.25 0.56 0.22 4.00 23a-806 45/1#26 as wt% 69.2 n.d. n.d. n.d. 0.6 n.d. 15.8 6.3 6.2 n.d. n.d. 98.1 prop. 2.90 0.05 0.55 0.37 0.37 4.00 23a-808 45/1#20 as-pb-sb wt% 68.5 n.d. n.d. n.d. 0.6 n.d. 15.7 4.7 6.4 2.6 n.d. 98.1 prop. 2.84 0.05 0.54 0.10 0.38 0.09 4.00 multiple ave. of 63 grains wt% 67.9 0.0 0.2 0.5 0.5 0.9 25.8 2.1 1.1 99.2 prop. 2.86 0.00 0.01 0.04 0.04 0.03 0.91 0.05 0.07 4.00 sd 2.77 0.13 0.84 1.03 0.42 1.53 5.23 2.35 2.36 kotulskite pd(te,pb,bi) 10-434 40/2#5 au wt% 40.4 n.d. 6.2 n.d. n.d. n.d. 52.5 n.d. n.d. n.d. n.d. 99.1 prop. 0.92 0.08 1.00 2.00 23a-806 75/1#6 pb wt% 42.8 n.d. n.d. 0.4 1.1 n.d. 52.0 3.1 n.d. n.d. n.d. 99.4 prop. 0.95 0.01 0.05 0.96 0.04 2.00 23a-798 63#2 bi wt% 42.5 n.d. n.d. n.d. n.d. n.d. 48.2 n.d. n.d. n.d. 9.1 99.8 prop. 1.00 0.92 0.08 2.00 multiple ave. of 13 grains   wt% 43.6 n.d. 1.2 0.1 0.2 n.d. 53.1 0.2 n.d. n.d. 1.2 99.6 prop. 0.97 0.01 0.00 0.01 0.99 0.00 0.01 2.00 sd 1.76 2.25 0.12 0.39 2.31 0.80 2.92 a close to empirical formula pd3-xte. ave.: average. prop.: atomic proportions. sd: standard deviation. n.d.: not detected. indicates no data. table 18 compositional types of atokite and cabriite sample no. grain no. type pd pt au cu fe sn te pb as sb sum 10-445 40/2#1 ‘pure’ a wt% 72.9 n.d. n.d. n.d. n.d. 26.7 n.d. n.d. n.d. n.d. 99.6 prop. 3.01 0.99 4.00 10-443 40#33 pt-as wt% 61.7 16.6 n.d. n.d. n.d. 14.2 n.d. n.d. 7.3 n.d. 99.8 prop. 2.63 0.39 0.54 0.44 4.00 10-443 40#11 au-pb wt% 66.2 2.1 3.8 n.d. n.d. 24.1 n.d. 3.3 n.d. n.d. 99.5 prop. 2.86 0.05 0.09 0.93 0.07 4.00 23a-806 45/2#25 au-as-te wt% 69.2 n.d. 2.6 n.d. 0.5 14.5 3.0 n.d. 6.9 2.1 99.2 prop. 2.80 0.06 0.04 0.53 0.10 0.40 0.07 4.00 10-443 40/2 pb wt% 65.8 n.d. n.d. n.d. 0.7 13.1 n.d. 19.8 n.d. n.d. 99.4 prop. 2.96 0.06 0.53 0.46 4.00 multiple ave. of 14 grains wt% 68.1 2.8 0.8 0.4 0.2 21.2 0.2 3.9 1.6 0.2 99.3 prop. 2.88 0.06 0.02 0.05 0.02 0.80 0.01 0.09 0.09 0.01 0.99 sd 3.58 4.81 1.38 0.64 0.36 5.49 0.80 6.91 2.90 0.56 multiple ave. of 5 grains   wt% 53.7 n.d. n.d. 17.5 0.7 27.3 0.3 0.2 n.d. n.d. 99.7 prop. 1.96 1.08 0.05 0.90 0.01 4.00 sd 1.15 0.71 0.45 2.76 0.58 0.40 a close to empirical formula pd3sn. ave.: average. prop.: atomic proportions. sd: standard deviation. n.d.: not detected. indicates no data. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 45 of 66 www.geusbul let in.org table 19 compositions of naldrettite, sopcheite, froodite, merenskyite, hessite and unnamed pd2te, pd2cutebi, (pd,ag)2te sample no. grain no.   pd au ag cu fe ni sn te sb bi pb s sum naldrettite pd2(sb,sn) 24-1034 40#115 wt% 59.2 3.4 n.d n.d 0.6 n.d 16.0 1.1 18.4 n.d n.d n.d 98.7 prop. 1.90 0.06 0.04 0.46 0.03 0.51 3.00 sopcheite (ag,cu)4pd3te4 24-1024 80#2 wt% 25.7 n.d 26.9 4.6 n.d n.d 1.8 36.7 n.d n.d 3.7 n.d 99.5 prop. 3.00 3.11 0.90 0.19 3.58 0.22 11.00 froodite pdbi2 24-1034 40#115 wt% 19.3 n.d n.d n.d n.d n.d n.d n.d n.d 79.7 n.d n.d 99.0 prop. 0.97 2.03 3.00 merenskyite pdte2 bs0304 45#203 wt% 28.3 n.d n.d 0.6 n.d n.d n.d 72.1 n.d n.d n.d n.d 101.0 prop. 0.95 0.03 2.02 3.00 hessite ag2te 10-434 40#5 wt% n.d n.d 61.7 n.d n.d n.d n.d 37.8 n.d n.d n.d n.d 99.5 prop. 1.98 1.02 3.00 unnamed pd2te 24-1048 40#24 wt% 67.1 n.d n.d 0.4 n.d n.d n.d 31.8 n.d n.d n.d n.d 99.3 prop. 1.97 0.02 1.97 3.00 unnamed pd2cutebi 18-978 40#29 wt% 34.2 n.d n.d 7.4 1.7 n.d n.d 20.6 n.d 35.7 n.d n.d 99.6 prop. 2.01 0.73 0.19 1.01 1.07 5.00 unnamed (pd,ag)2te 24-1022 63#14 wt% 47.3 1.7 10.8 1.3 0.6 n.d n.d 34.6 n.d n.d 3.3 n.d 99.6 prop. 1.53 0.03 0.35 0.07 0.04 0.93 0.05 3.00 prop.: atomic proportions. n.d.: not detected. indicates no data. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 46 of 66 www.geusbul let in.org fig. 32 2 columns a b c d e f g h i j k 23a-798, 40-1 #18 10 μm 23a-798, 40-1 #128 10 μm 23a-808, 45-1-1 #1 10 μm 23a-807, 45-1 #12 10 μm 23a-806, 45-1 #19 10 μm 10-443, 40-2 #19 10 μm 23a-807, 45-1 #1 10 μm 23a-808, 75 #22 10 μm 23a-808, 45-1-1 #12 10 μm 10-445, 40-2 #1 10 μm 22-977, 40 #1 10 μm fig. 32 bse images of pge arsenides in polished monolayer thin sections of hs concentrates. as for the intermetallic compounds (see text and fig. 31), arsenides are related to a wide range of precious metal phases and sulphides, while local geochemical balances control the parageneses. table 3 provides formulae for the precious metal minerals. formulae for common rock-forming minerals and base metal sulphides are found in appendix 1. formulae are given here for rare minerals or compositional varieties that are not listed in table 3 or appendix 1. abbreviated names of minerals, compounds and alloys are as follows: apd: arsenopalladinite/stillwaterite (pd8as3). kt: kotulskite. sp: sperrylite. vnc: vincentite. gng: guanglinite (obsolete and renamed to vincentite). plas: palladoarsenide. ism: isomertiete (pd11as2(sb,sn)2). mjk: majakite. zv: zvyagintsevite. at: atokite. pd11as2sn2: unnamed compound. aucu: tetra-auricupride. (au,cu,ag): alloys dominated by au and cu. (au,ag): native gold. bn: bornite. chlc: chalcocite. hb: hornblende. act: actinolite. mt: magnetite. grain numbers and scales included in the images. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 47 of 66 www.geusbul let in.org table 20 compositional types of arsenopalladinite and palladoarsenite sample no. grain no. type   pd pt au cu fe sn te pb as sb sum arsenopalladinite pd8as3 10-434 40#8 ’pure’ a wt% 77.0 n.d. n.d. 0.7 0.7 n.d. n.d. n.d. 21.0 n.d. 99.4 prop. 7.75 0.12 0.13 3.00 11.00 23a-798 40#2 au-sn wt% 69.7 n.d. 6.9 1.4 n.d. 5.1 n.d. n.d. 16.8 n.d. 99.9 prop. 7.35 0.39 0.25 0.49 2.51 11.00 23a-806 45/1#11 cu wt% 74.7 n.d. n.d. 3.4 0.3 n.d. n.d. n.d. 21.4 n.d. 99.8 prop. 7.37 0.56 0.06 3.01 11.00 23a-807 45/2#35 te-sn-sb wt% 71.0 n.d. n.d. 1.7 0.7 6.2 8.9 n.d. 9.1 2.2 99.8 prop. 7.58 0.30 0.15 0.59 0.80 1.38 0.20 11.00 multiple ave. of 67 grains wt% 73.6 n.d. 1.7 1.7 0.6 2.0 0.2 n.d. 19.6 0.1 99.4 prop. 7.48 0.09 0.29 0.11 0.18 0.02 2.83 0.01 11.00 sd 2.90 2.26 1.11 0.45 0.15 2.15 2.20 0.10 palladoarsenite pd2as 18-972 80#2 ‘pure’ b wt% 71.8 n.d. n.d. 1.2 0.4 n.d. n.d. n.d. 26.0 n.d. 99.4 prop. 1.93 0.05 0.02 0.99 3.00 10-434 40/2#1 pt wt% 59.7 12.4 n.d. 1.8 1.0 n.d. n.d. n.d. 24.5 n.d. 99.4 prop. 1.69 0.19 0.08 0.05 0.98 3.00 23a-808 40/1#17 cu-fe wt% 65.8 n.d. n.d. 4.6 2.3 n.d. n.d. n.d. 27.0 n.d. 99.7 prop. 1.70 0.20 0.11 0.99 3.00 multiple ave. of 6 grains   wt% 67.2 3.9 n.d. 1.8 0.9 0.6 n.d. 0.4 24.7 n.d. 99.5 prop. 1.84 0.06 0.08 0.04 0.02 0.01 0.96 3.00 sd 6.32 6.03 1.78 0.73 1.47 0.90 1.68 a close to empirical formula pd8as3. b close to empirical formula pd2as. prop.: atomic proportions. sd: standard deviation. n.d.: not detected. indicates no data. table 21 compositional types of vincentite/guanglinite and compositions of isomertieite, majakite and unnamed pd11as2sn2 sample no. grain no. type pd pt au cu ni fe sn te pb as sb sum vincentite/guanglinite pd3as 10-445 40/1#2 ‘pure’ a wt% 81.5 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. 18.8 n.d. 100.3 prop. 3.01 0.99 4.00 10-443 40#33 pt wt% 74.7 4.6 n.d. n.d. n.d. 0.6 3.1 n.d. n.d. 16.4 n.d. 99.4 prop. 2.86 0.10 0.04 0.11 0.89 4.00 23a-806 45/1#7 au-sn wt% 69.7 n.d. 6.7 1.0 0.6 4.5 n.d. n.d. 16.1 n.d. 98.6 prop. 2.71 0.14 0.07 0.04 0.16 0.89 4.00 18-988 40#7 cu-te-pb wt% 68.2 n.d. n.d. 4.1 n.d. 0.7 n.d. 11.8 6.4 8.1 n.d. 99.3 prop. 2.70 0.27 0.05 0.39 0.13 0.45 4.00 10-443 70#4 sn wt% 78.4 n.d. n.d. n.d. n.d. n.d. 7.7 n.d. n.d. 13.6 n.d. 99.7 prop. 3.00 0.26 0.74 4.00 23a-808 45/1#51 pb wt% 70.3 n.d. n.d. n.d. n.d. 0.4 n.d. 2.8 19.1 7.2 n.d. 99.8 prop. 3.01 0.03 0.10 0.42 0.44 4.00 23a-808 45/1#18 sb wt% 75.2 n.d. n.d. 1.0 n.d. 0.3 n.d. n.d. n.d. 12.1 11.0 99.6 prop. 2.88 0.07 0.02 0.66 0.37 4.00 multiple ave. of 64 grains wt% 76.0 0.3 0.3 0.6 n.d. 0.5 2.6 1.7 1.6 14.9 0.4 98.8 prop. 2.91 0.01 0.01 0.04 0.04 0.09 0.06 0.04 0.80 0.01 4.00 sd 4.12 1.45 2.10 1.26 0.48 2.63 4.73 6.07 4.62 3.47 isomertieite (pd,cu)11as2(sb,sn)2 multiple ave. of 15 grains wt% 71.4 0.3 0.4 1.8 n.d. 1.0 1.1 1.2 n.d. 9.5 12.6 99.4 prop. 10.38 0.02 0.02 0.45 0.27 0.13 0.15 1.97 1.61 15.00 sd 1.79 0.67 0.58 0.91 0.97 1.79 1.93 0.49 2.42 majakite pdnias 22-977 125#1 wt% 44.3 n.d. n.d. n.d. 24.5 0.7 n.d. n.d. n.d. 30.5 n.d. 100.0 prop. 1.00 1.00 0.03 0.98 3.00 unnamed pd11as2sn2 10-445 40/2#1 wt% 75.3 n.d. n.d. n.d. n.d. n.d. 14.6 n.d. n.d. 9.5 n.d. 99.0 prop. 11.09 1.93 1.99 15.00 a close to empirical formula pd3 as. prop.: atomic proportions. ave.: average. sd: standard deviation. n.d.: not detected. indicates no data. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 48 of 66 www.geusbul let in.org 5 discussion: constraints for the mineralisation model orthomagmatic origins were advocated for the skaergaard pge-au mineralisation by andersen et al. (1998), nielsen et al. (2005), holwell & keays (2014), holwell et al. (2015, 2016) and keays & tegner (2016). the proposed models assume fractionation and enrichment in precious metals, cu and s until sulphide saturation was reached and led to gravitational accumulation of sulphide melt with precious metals in the intermittent floor of the magma chamber. the sulphide – precious metal paragenesis subsequently re-equilibrated during cooling. existing models are, among others, inspired by the merensky reef of the bushveld complex (naldrett 2004). a more complex but nevertheless comparable gravitational model involves a first stage of crystallisation and fractionation below the roof of the magma chamber, scavenging of precious metals by immiscible sulphide melt in crystal mush, followed by dissolution in dense fe-rich immiscible silicate melts and slumping to the floor (nielsen et al. 2015). the difference between this and the classic model is that the first stage of sulphide saturation occurs in evolved mush melt below the roof of the intrusion rather than in contemporary and lesser evolved bulk melt. in the floor mush, gravitational sorting resulted in macrorhythmic layers with stratigraphic concentrations of mush melt in which fractionation once again led to sulphide saturation and concentration of precious metals and subsequently to reaction with residual and immiscible fe-rich mush melt. the reactions led to very high precious metal to cu ratios and transport of precious metals in upwardly-migrating mush melts (nielsen et al. 2015). a third model (e.g. wernette et al. 2020) argued for scavenging of precious metals by upwardly-migrating fluids and possibly residual silicate melt (boudreau 2004) from precious metal–bearing interstitial sulphides deep in already crystallised gabbros of the intrusion. this later model was developed for the j-m reef in the stillwater igneous complex (boudreau et al. 1986). the models are very different, but all revolve around the timing of crystallisation processes relative to the loss of residual mush melts, the timing of sulphide saturation, the permeability of the gabbroic mush and the role of free volatile phases for transport and deposition of precious metals, especially au. the three model types can be simplified to a downward accumulation model, a dissolution and upward accumulation model, and the model by nielsen et al. (2015) that combines the two contrasting models into two successive stages. a first concentration of precious metals in the floor was caused by slumping of enriched crystal mushes that originated from under the roof of the magma chamber. this was followed by a second stage of fractionation and deposition of precious metals in upwardly-migrating mush melts of the macrorhythmic layers. in the following sections, we constrain the precious metal composition of the remaining bulk liquid at sulphide saturation, the temperatures and the relative timing of the processes that led to the now observed mineralisation. the discussion is based on an integration of the stratigraphic and temporal relationships presented in nielsen et al. (2015, 2019a), and the summaries of petrography and compositions of sulphides and precious metal phases of the skaergaard pge-au mineralisation presented here. 5.1 concentrations of pge and au in the magma at sulphide saturation a high feo total (feo*) in the remaining bulk liquid ensured that precious metals were retained until sulphide saturation. the formation of the skaergaard pge-au mineralisation was delayed until upper mz crystalised, when three quarters of the bulk magma had solidified (nielsen 2004; andersen 2006). exceptions are observed halfway up mz due to local sulphur saturation in mushes and in reaction with rafts of roof rocks or other lithic inclusions (e.g. bird et al. 1991; nielsen et al. 2015). the composition of plagioclases in equilibrium with the sulphides of the mineralisation corresponds to plagioclase in uza gabbros, when 18% (f-value 0.18) of the bulk liquid remained (nielsen et al. 2019a). the macrorhythmic layers that host the mineralisation in the upper mz, however, formed when 27% (f-value 0.27) of the volume of the bulk magma remained (nielsen et al. 2015). this discrepancy supports that sulphide saturation occurred in the crystal mush. the mush formed below the roof of the intrusion, and was enriched in pges, au, dense phases (pyroxene, magnetite and ilmenite) and immiscible fe-rich silicate melt (formed at f = 0.1; nielsen et al. 2015). convection carried the dense mush to the floor and built the succession of macrorhythmic layers that now host the mineralisation. this crystallisation and mineralisation model was fully developed by nielsen et al. (2015, 2019a), while nielsen et al. (2019b) modelled the corresponding evolution of the bulk liquid in the intrusion. in upper mz, sulphide saturation in the mushes of proto-macrorhythmic layers would therefore have taken place after crystallisation of approximately one third of the bulk composition of material in the macro-layers, while f decreased from 0.27 to 0.18. the residual melt https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 49 of 66 www.geusbul let in.org in the crystal mush of the proto-macrorhythmic layers would, due to the density contrasts between solids and melts (i.e. plagioclase floats in the silicate melt; nielsen et al. 2019b), be concentrated in the lower part of the plagioclase layers that forms the tops of the proto-macrorhythmic layers (nielsen et al. 2015). ideally, and assuming no loss of cu or precious metals, the remaining bulk liquid would at this stage have feo* c. 22 wt%, pt 10 ppb, pd 63 ppb, and au 42 ppb at 27% remaining bulk liquid (nielsen et al. 2019b). the mush melt would reach sulphide saturation after crystallisation of one third of its volume and ideally contain pt c. 15 ppb, pd 94 ppb and au 63 ppb. these elemental concentrations apply only to the melts in mushes below the roof of the intrusions from where precious metals were fed to the floor (nielsen et al. 2015). the relative proportions between crystal mush residing at the intermittent floor and new mush descending from the roof of the magma chamber are not constrained. it is therefore at present not possible to estimate the compositions of mush melt and the thickness of the mush at the floor. thicknesses of mush between a few metres and >100 m are estimated in holness et al. (2017) and apparently vary across the intermittent floor of the magma chamber. 5.2 constraints on the timing and temperature of sulphide saturation the mineralisation is hosted in gabbros in the upper 100 m of mz. based on the stratigraphic height in the ls and following thy et al. (2013), these gabbros would have crystallised at temperatures between 1030 and 1050ºс. this temperature estimate is based on the correlation between the composition of liquidus plagioclase and temperatures on the lld (liquid line of descent) of the bulk liquid (thy et al. 2013). another constraint on the relative timing of the mineralisation is the correlation between pd5, the main pge mineralisation level, gabbros rich in ilmenite and ti-magnetite and a high plagioclase to clinopyroxene ratio (plag/px) of c. 2 (nielsen et al. 2015). this concentration of the liquidus phases with the highest and lowest densities in the one and the same concordant cumulus layer is at first enigmatic. nielsen et al. (2015) suggested that plagioclase is on the liquidus in the contemporary bulk liquid, and that fe and ti-rich paragenesis crystallised from immiscible fe-rich interstitial silicate melt left behind in the layer. crystallisation from contrasting melt compositions is evidenced by resorption and replacements of plagioclase by symplectites at contacts with anhedral feti oxides (fig. 6 in nielsen at al. 2015). this is supported by the presence of large euhedral apatite crystals (in fig. 5e in nielsen et al. 2015) inside anhedral magnetite, reaction rims and symplectites adjacent to the liquidus clinopyroxene and resorbed plagioclase (see fig. 6 in nielsen et  al. 2015). these petrographic relations show that the anhedral magnetites crystallised from highly evolved mush melts comparable to the bulk liquid at the base of uzb (f = 0.1). the upward density-controlled layering in precious metal–rich macrorhythmic layers (nielsen et al. 2015) includes from base to top: (1) melanocratic gabbro with increasing plag/px, with or without accumulations of feti liquidus crystals; (2) average gabbro; (3) a zone that is high in plag/px, rich in interstitial feti oxides, and host to pd concentrations, for example, pd5 in ml0; (4) a plagioclase-rich top grading into a high plag/px gabbro that is poor in feti oxides at the transition to the overlying macrorhythmic layer (nielsen et al. 2015). the density stratification is stable, assuming that the plagioclase-rich gabbros with high proportions of feti oxide and high plag/px are composed of crystals of liquidus plagioclase and pyroxene that are out of equilibrium with the interstitial fe-rich silicate melt from which the feti oxides crystallised (nielsen et al. 2019a). the base of uzb and the ponding of immiscible fe-rich melts (nielsen et al. 2019b) occurred at c. 1015°с (thy et al. 2013). the reaction between the already formed sulphide droplets and interstitial fe-rich melt in the mushes of the proto-mineralisation would then, assuming the same liquidus temperatures in mush melt, have occurred at c. 1015ºc. the temperature interval for sulphide saturation and formation of droplets of sulphide melt would be 35°c at most and may be as little as 15°c. 5.3 petrographic evidence for an initial high-t origin of the mineralisation a first orthomagmatic stage in the development of the skaergaard cu-fe sulphides and pge-au mineralisation is evidenced by sulphide droplets hosted in primary ilmenite and pyroxenes of the gabbroic host (nielsen et al. 2005; godel et al. 2014; holwell et al. 2015; nielsen et al. 2015; rudashevsky et al. 2015). numerous droplets of fe-cu sulphides have been imaged and analysed (see fig. 7g–n), some with shapes of negative crystals indicating trapping during the growth of the host (nielsen et al. 2015). the droplets often contain precious metal minerals and alloys, such as (pd,cu) alloys (fig. 18), pd sulphides, (au,cu) alloys (fig. 14a, b) and (cu,au) alloys (fig. 24a–j). the presence of droplets of sulphide melt with precious metal grains and crystals was further demonstrated in the hrxct study of godel et al. (2014). these sulphide droplets in liquidus phases of the gabbroic host are up to several hundred microns across (fig. 15), have a sphericity >0.9 and contain grains and crystals of https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 50 of 66 www.geusbul let in.org skaergaardite (pdcu) with an average grain size of 19.2 µm from a total of 25 sulphide grains. the wide distribution of droplets of cu-fe sulphide melts and precious metals (figs 16a–l, 17a–k, m) is exceptional. these polyphase grains of cu-fe sulphide, cu-fe sulphide with grains of (cu,pd) alloys and of (cu,pd) alloys occur in primary rock-forming minerals (figs 6d–g, 7a, b, d) as well as between rock-forming minerals (figs 6b, c, 7c, e). similar compositions of (cu,pd)β and (cu,pd)α alloys are found in free droplets, in intergrowths with skaergaardite (cupd) and nielsenite (cu3pd; fig. 17), as well as in grains related to cu-fe sulphides. the occurrence of isolated droplets of (сu,pd)α alloy and chalcocite in the ilmenite host is critical (fig. 6f). it demonstrates the formation and presence of more types of sulphide droplets during the crystallisation of the rock-forming minerals. the ratio of sulphide to precious and other metals in droplets and globules varies widely (fig. 16a–l) and reflects the lack of simple equilibrium with a common bulk silicate liquid. monomineralic microdroplets of (pd,cu)α are also observed in host-rock ilmenite (fig. 6e) and in (pd,cu) alloys within cu-fe sulphide droplets in all the precious metal–bearing mineralisation levels (figs 16c, 17c). the compositional differences between the cu-pd dominated pd5 in the lowest and central parts of the mineralisation (table 7) and the cu-au dominated upper au-rich pd2-pd1/au mineralisation levels (table 11), agree with the decrease in distribution coefficients (d-values) from pd to au between sulphide and silicate melts. consequently, the difference between the pd5 and pd2-pd1/au levels is consistent with high-temperature, and the magmatic nature of the first-formed sulphide-precious metal paragenesis. the formation of (cu,pd) droplets is a consequence of a cu-rich and s-poor system with variable ratios of cu to precious metals (e.g. karup-møller et al. 2008). high skaergaardite (pdcu) would crystallise in equilibrium with bornite and a precious metal–rich and very low-s melt with a composition close to that of skaergaardite at t ≥ 900°c (karup-møller & makovicky 1999). the solubility of precious metals in cu-rich sulphide melts is generally too high to allow nucleation of (cu,pd) and (cu,au) compounds (liu & brenan 2015) and the skaergaard (cu,pd) compounds may have formed from unprotected sulphide droplets subjected to resorption and loss of cu and s, caused by reaction with interstitial fe-rich silicate melt in the crystal mush (nielsen et al. 2015). primary pd sulphides, including vysotskite and vasilite, formed at temperatures lower than those of their cu-pd microglobule hosts (fig. 19b, c). the pd sulphides probably formed by exsolution from the sulphide melt, like the globules themselves. constraints on the relative timing of the formation of the precious metal phases and parageneses are provided by the common observation of ‘droplets within a droplet’. for example, droplets of vysotskite and vasilite in bornite, or droplets of pentlandite in chalcopyrite (fig. 28), as well as grain zonation and exsolution textures. the almost perfect vysotskite-bearing chalcocite-bornite droplet enclosed in primary ilmenite in fig. 28a illustrates the trapping of a droplet of cu-rich and fe-poor sulphide melt while the ilmenite host crystallised from the silicate liquid in the mush of mineralisation level pd2a in drill core 90-18. sulphide saturation was reached before the ilmenite grain crystallised in the mush melt of the macrorhythmic layer, and before the remaining bulk liquid had reached sulphide saturation. figure 28b illustrates the formation of a pgerich globule within a bornite grain. it is interpreted to be the result of dissolution and loss of s, cu and fe from firstformed sulphide droplets causing enrichment in precious metals and subsequently to exsolution of an immiscible pge-rich melt from the cu-rich sulphide melt (see karupmøller & makovicky (1999) and karup-møller et al. (2008) for supporting experimental evidence). the relationships between droplet-like grains of (au,cu) alloy (fig. 24) and cu-fe sulphide hosts in which such grains are found (fig. 23b–e), as well as the unique relationships between sulphide and cu-au metallic melts (fig. 23b–f, i–k) are similar to those observed for (cu,pd) alloys and indicate an initial high-t magmatic origin of the observed (au,cu) phases. further support is provided by the coexistence of ordered phases like auricupride (au3cu), tetra-auricupride (aucu), unnamed aucu3, and non-stoichiometric alloys of (au,cu) and (cu,au), some of which are zoned (fig. 25a–g; table 12) and intergrown (fig. 25h–j; table 13). although they are affected by later low-t exsolutions, they represent the original high-t grains or droplets of (cu,au) alloys. additionally, cu-fe sulphides form interstitial aggregates with inclusions of (cu,au) alloys between the minerals of the gabbroic host – like those described above – as well as individual grains of the same (au,cu) phases (figs 23a, 27a–j). the same paragenesis is found throughout the skaergaard pge-au mineralisation, although with highly variable relative proportions of cu-fe sulphides, precious metals and groups of precious metal phases. the diversity of the parageneses and the high variability in the compositions of the precious metal phases negate that they were in equilibrium with a common bulk liquid. local geochemical conditions controlled the compositions of the preserved paragenesis. but all parageneses formed from parental immiscible high-t sulphide melts that evolved in parallel during cooling and crystallisation of interstitial melts of mineralisation levels in the crystal mush of the intermittent floor of the intrusion. the liquidus temperatures of non-stoichiometric (au,cu,pd) and (au,cu,ag) alloys (fig. 33a, b) are estimated to be 1200–1000°c and 1050–970°c, respectively. in these three-phase diagrams, the lowest liquidus temperatures of the copper-dominated (cu,au,ag) alloys with https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 51 of 66 www.geusbul let in.org little pd are <1100°c (labelled “6” in fig. 33a) compared to the au-dominated (au,cu,pd) alloys with up to 1200°c (labelled “5” in fig. 33a). although the temperatures in fig. 33 for (pd,cu) and (ag,cu,au) alloys are not accurate for more complex natural systems, they do support that these primary alloys formed at magmatic temperatures during the crystallisation of the gabbroic host. 5.4 the role of coexisting fluids in the skaergaard pge-au mineralisation 5.4.1 evidence for coexisting magmatic fluid the coexistence of hydrous silicate globules and globules of sulphide with precious metal phases in liquidus feti minerals (fig. 15) shows that the skaergaard pge-au mineralisation formed in the presence of coexisting silicate melt, sulphide melt and a free, volatile-rich phase (godel et al. 2014). the silicate globules are, for example, composed of tremolite, hornblende, cummingtonite, pargasite, biotite, talc, albite, quartz and magnetite. some of the globules contain au-phases (holwell et al. 2016). the hydrous silicate globules are in some cases very frequent and more than 30 have been found in a single feti oxide grain and represent a noticeable volume (fig. 15e). the silicate globules have a sphericity >0.95 and diameter 10–190 μm (average 57 μm; godel et al. 2014) and coexisted with the silicate melt from which their hosts crystallised. the coexistence of silicate and sulphide melts with a free volatile phase is also described for the norilsk-talnakh ni-cu-pge mining camp (barnes et al. 2019). 5.4.2 nature of the coexisting fluid phase gold occurs only in the atomic state au0 in silicate melt. due to its siderophile-chalcophile affinity, au is selectively concentrated in sulphide melt, feti oxide or metal fractions in the form of alloys with au, cu, pd, pt, ag and ni. however, au may exist in several forms in supercritical fluid at t < 1100°c, and, for example, as au(oh)0 and as au+ complexes at high temperatures in fe-rich systems with high fe2+ activity (e.g. spiridonov 2010a). the solubility of au and ag is high in coexisting, hydrous, magmatic fluids and leads to their upward transportation during fractional crystallisation. cl-containing fluids are traditionally seen as transporters of au at high temperatures in magmatic systems (e.g. spiridonov 2010a; zajacz et al. 2013). the presence of cl in the fluids of gabbros from near the margins of the intrusion is shown by the occurrence of cl-bearing hornblende with up to 2 wt% cl (sample 23a-808), cl-apatite and the copper chloride atacamite (cu2cl(oh)3; fig. 9r) in association with pges and bornite and cl-apatite in gabbros from the centre of the mineralisation. larsen et al. (1992) and larsen & brooks (1994) describe the very common occurrence of fluid inclusions related to lower-t pegmatites. they equilibrated at 770–655°c and at an oxygen fugacity (ƒo2) of 1.5 to 2 log units below the quartz-fayalite-magnetite (qfm) buffer, well below both the temperature (>1000°c) and the qfm buffer at which the mineralisation first formed (bollingberg 1995). all evidence for low ƒo2 conditions relates to feti oxide pairs that recrystallised and re-equilibrated at subsolidus conditions. none of these represent the ƒo2 as precious metals were first deposited in the gabbros. hanley et al. (2005) demonstrated that single-phase hypersaline fluids that interact with magmatic liquids and partially crystallised rocks can dissolve and transport significant amounts of au and pt. this may well a 1 2 4 6 5 1+3 15 00 º 14 00 º 13 00 º 1200º 1100º 1000º au cupd 1000º 10 50 º 950º 900º 850º 80 0º 95 0º 10 00 º b au (au,cu,ag) ag cu fig. 33 one column fig. 33 phase diagrams showing compositions (wt%) of a: primary cu, pd and au and b: primary alloys of au, cu and ag on experimental liquidus surfaces of the au-cu-pd system (°c; after savitsky (1984). compositional fields shown in (a) for the au-cu-pd system are as follows: 1: skaergaardite. 2: nielsenite (cu3pd). 3: (pd,cu)β alloys. 4: (cu,pd)α alloys. 5: (au,cu,pd) alloys. 6: (au,cu,ag) alloys. compositional field shown in (b) for the au-cu-ag system includes (au,cu,ag) alloys. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 52 of 66 www.geusbul let in.org explain the correlation between pt and au and low pd/ pt ratios in the upper au-rich mineralisation levels, as well as the high pd/pt and pd/au ratios in the lowermost pd6–pd4 mineralisation levels. the latter are caused by the high d-value for pd in sulphide melt, as well as loss of au and pt to upward-migrating residual melts and volatile-rich fluids (nielsen et al. 2015). such residual melts in equilibrium with high-t fluids are enriched in cl, h2o, te, se, s and hs–, and metals such as au, ag, pb, zn, pd, sn, sb and bi. they stabilise au complexes with h2o, cl, bi-sulphide and arsenide-sulphide compounds, which are all important for transport and ore formation in, for example, volcanogenic and plutonic au deposits. strong correlations between elemental pairs such as au-as, au-te, au-se, au-bi and au-sb, and the high se and te in gabbros immediately above the au-rich mineralisation levels (holwell & keays 2014; holwell et al. 2016; keays & tegner 2016) show that telluride and selenide complexes played a leading role in the migration of au-bearing hydrous residuals just as in epithermal volcanic au deposits (e.g. yakich et al. 2021). laser ablation inductively coupled plasma mass spectrometry (la-iсp-ms) analyses of cu-fe sulphide blebs from levels pd6 and pd5 show anomalous levels of small chalcophile elements, such as pd, au, ag, plus se (up to c. 650 ppm) and te (c. 100 ppm; holwell et al. 2015). the concentrations are much higher in the au-rich mineralisation levels, with se and te reaching 1000 and 1150  ppm, respectively. the high concentrations and the correlations between au, te and se support transport of au in solution in the form of hetero-polynuclear telluride, selenide and sulphide-arsenide complexes in fluids. this supports the models for transportation and deposition of gold presented in rudashevsky et al. (2014), nielsen et al. (2015) and wernette et al. (2020). the occurrence and increasing importance of stable metalloid-complexes in the coexisting fluid phase during the formation of the au-rich mineralisation levels is reflected in the contrasting compositions of skaergaardite (pdcu). the correlation between au and te in skaergaardite is +0.38 at the 95% confidence level (table  5, 1482 analyses in 24 samples) but cannot be crystallochemical as the two elements have different positions in the structure of skaergaardite. a correlation between au and te is supported by the factor analysis based on 311 analyses of skaergaardite (pdcu) from pd5 sample 24-2057 (fig. 12). the first factor (f1) is au-te-cu and represents the main variation in precious metal paragenesis up through the mineralisation levels from pd6 to pd1/au and au+1 (rudashevsky et al. 2004, 2014). the correlations and anticorrelations in these levels corroborate with the factor analysis based on bulk-rock compositions in nielsen et al. (2019a). 5.4.3 extent of equilibration with aqueous fluids hydrous silicates are very rare in gabbros in the central parts of the intrusion, except for biotite related to feti oxides (holness et al. 2011), although not all residual melt and fluid escaped from the depths of the intrusion (wernette et al. 2020) or from the mineralised levels. in stark contrast, gabbros closer to the margins (drill cores 90-10 and 90-23a) have hydrous silicate in their matrixes and preserved “paired inclusions” of si-rich and fe-rich melt (holness et al. 2011). they also contain significantly higher proportions of intermetallic precious metal compounds with metalloids and fluid-related elements such as as, pb, te and s. lateral contrasts are also seen in the proportions of teand as-bearing minerals in the au-rich mineralisation levels (figs 10, 26a–e). the proportions are lower in the central parts of the mineralisation compared to gabbros nearer the margins of the intrusion. they host the characteristic association of (au,cu,ag) alloys with platinum group arsenides and tellurides (fig. 26g–j), e.g. in the au-rich mineralisation level pd4 in drill cores 90-10 and 90-23a (nielsen et al. 2005, 2015, 2019a; rudashevsky et al. 2014). only a few grains of au minerals, such as tetra-auricupride (aucu), and te minerals, such as keithconnite (pd3-xte), can be found in the central parts of the mineralisation. here, the coexisting fluid that carried volatile components and metalloids was seemingly lost from the mushes. the contrast between central areas and those close to the margins of the intrusion led nielsen et al. (2015) to suggest trapping of residual silicate melts and a free magmatic volatile phase in the faster-cooled gabbros nearer to the margins. this, however, does not exclude inflow of meteoric water (e.g. norton & taylor 1979) through contraction cracks in already crystallised gabbros (norton et al. 1984; bird et al. 1986, 1991). these fractures were filled by secondary ca-rich clinopyroxene amphiboles ranging from actinolite to hornblende, biotite, feti oxides and calcic plagioclase (manning & bird 1986). the tendency for larger grains of precious metal phases at tof may reflect contrasts in the crystallisation environment between the centre and the margins of the intrusion. in addition, and possibly due to loss of permeability during cooling near the walls of the magma chamber, trapped residual melt and fluids formed late veins with small grains of precious metal phases. 5.5 low-t re-equilibration cu-fe sulphides, some with inclusions of precious metal minerals, formed from primary sulphide melts during crystallisation of the gabbroic host. the cu-fe sulphides display solid state exsolution textures in response to cooling (see fig. 5k–n). the temperature of the cu-rich https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 53 of 66 www.geusbul let in.org sulphide (iss) exsolution is 490–350°с (kullerud 1957, 1967; spiridonov 2010b). an important temperature constraint is, however, provided by orthopyroxene-clinopyroxene pairs (e.g. fig. 7c). hot, primary droplets of the cu-fe sulphide melt in the margins of augites slowed the cooling of augite and exsolution of orthopyroxene. we calculated equilibrium temperatures for pyroxene pairs in samples 24-1057 and 18-1010 from pd5 using the two-pyroxene thermobarometer of brey & köhler (1990). the estimated temperatures are 808–829°с (р = 1 kbar) and 817–839°с (р = 4 kbar) and constrain droplet formation to temperatures well above 800°c. the experimental solidi for the cu-pd system are 1554–1083°с and for cu-au 1063–1083°c at increasing cu (figs 34, 35; baker et  al. 1992) with continuous, structurally disordered series of cubic cu-pd and cu-au solid solutions. during slow cooling and equilibration, the cu-pd and cu-au compounds obtain ordered structures as exemplified by ‘microdroplets’ of (cu,pd) and (cu,au) alloys with sub-euhedral to euhedral habits in sulphide hosts (figs 16m–p, 17l, 23g, h). these euhedral crystals of precious metal phases may represent droplets of precious metal melt that re-equilibrated and recrystallised at lower temperatures. below 871 k, equivalent to 598°c, and in the compositional range 0.43 < cu < 0.70 at.%, the pd-cu phase compares to cubic cscl (space group pm3m). compositions with 0.53 < cu < 0.65 correspond to a single ordered phase, and in the intervals 0.43 < cu < 0.53 and 0.65 < cu < 0.70 they correspond to a single ordered phase coexisting with a disordered solid solution phase (fig.  34a; subramanian & laughlin 1991). the ordered stoichiometric intermetallic pdcu compound is skaergaardite (rudashevsky et al. 2004) and the two areas of partial ordering of the solid solution seem to correspond to the non-stoichiometric (cu,pd)β compound. ordering of pdcu3-type compositions occurs at t  <  508°c (fig. 34b; subramanian & laughlin 1991). at 20–25 at.% pd (fig. 34b), the phase has a tetragonal structure (space group p4mm) and corresponds to pdcu3 compound, nielsenite (pdcu3; мcdonald et al. 2008). in the interval 10–20 at.% pd, it has a cubic structure similar to that of aucu3 (space group рm3m). this phase seems to correspond to the non-stoichiometric compound (cu,pd)α. the degree of cu and pd ordering decreases from skaergaardite (pdcu) to nielsenite (pdcu3) in parallel with the increasing disorder in pdcu3 with an excess of pd (schubert et al. 1955). nielsenite (pdcu3) therefore forms at lower temperatures than skaergaardite (pdcu), which is also suggested by the occurrence of skaergaardite rimmed by nielsenite (e.g. fig. 17g). the mineralisation also exemplifies the highto low-t re-equilibration of cu-au compositions of tetra-auricupride (aucu), to auricupride (cu3au) and unnamed au3cu. grains of cu-au minerals are zoned (fig. 25a–g; table 12) and intergrowths of these minerals illustrate exsolution from parental high-t compositions (fig. 25h–j; table 13). the au-rich phases formed in the following order: (au,cu,pd) or (cu,au,pd) alloys, followed by ordered and stoichiometric compositions of aucu or aucu3 and silver-containing solid solutions at lowest temperature. ≈≈ β sk β nls 300 cu 20 40 pd in cu+pd (at.%) 60 80 pd 500 700 900 1100 1083 b 1554 1300 t (º c ) t (k ) α 500 600 700 800 900 50 sk nls pd 60 70 cu cu in cu+pd (at.%) csci 871 k fcc + csci fcc + csci fcc-structural type solid solution structural type a fig. 34 one column fig. 34 binary phase diagrams for cu and pd after savitsky (1984) and baker et al. (1992). a: stability fields for structural types cscl and for fcc + cscl for composition with 0.43 < cu < 0.70 (at.%). abbreviations in (a) are as follows: fcc: face centered cubic. sk: skaergaardite. nls: nielsenite. b: stability fields for coumpounds and minerals from au to cu. solid lines: established boundaries of stability fields. dashed lines: approximate positions of the boundaries. abbreviations in (b) are as follows: nls: nielsenite (cu3pd). sk: skaergaardite. α: (cu,pd)α alloy. β: (pd,cu)β alloy. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 54 of 66 www.geusbul let in.org the complex binodes of the low-t solid solution in the au-cu phase diagram correspond to ordered phases of aucu, au3cu, and aucu3 (fig. 35; okamoto et al. 1987; baker et al. 1992). the ordering of aucu, au3cu and aucu3 occurs at t < 410–380°c. aucu has an orthorhombic structure at t < 410–380°c (fig. 35) and has a tetragonal structure at t < 380°c. aucu3 alloy is cubic and analogous to auricupride at t < 390°с (fig. 35). in addition, aucu, auricupride (aucu3), and unnamed mineral au3cu also formed as the result of exsolutions of au-cu solid solutions. phases of cu-ag appear to preserve comparatively high-t origins. rare grains of cu and ag with 5.8 wt% cu (table 3, row 1, table 9, row 8), and appear to have formed at 720°c (savitsky 1984). the spatial relationship between intermetallic compounds of cu and pd (skaergaardite; pdcu) and intermetallic compounds of au and cu (tetra-auricupride; aucu) in fig. 19i also reflects decreasing temperatures. this is illustrated by a skaergaardite composition that equilibrated to ‘low-skaergaardite’ at c. 600°c (subramanian & laughlin 1991; karup-møller et al. 2008) and which was subsequently rimmed by tetra-auricupride (aucu) that equilibrated at c. 380°c (okamoto et al. 1987). 5.6 petrographic information relevant to beneficiation of the ore grain size and mineralogical associations of the precious metal phases are important constraints for beneficiation of the mineralised gabbros and potential ores of the skaergaard pge-au mineralisation. the bse images of the 30 samples presented in this study suggest a range of grain sizes from 1 to 90 µm (measured as ecd; see tables 1 and 3). the average grain sizes of precious metal phases in samples from the central intrusion (tof), the eastern margin (core 90-23a), and the western margin (drill core 90-10) are 23.6 ± 0.5, 19.3 ± 0.8 and 22.5 ± 1.5 µm, respectively (fig. 36). the precious metal grains are marginally larger in more central parts of the intrusion. heavy mineral hs concentrates also show that precious metal minerals in the –125 μm fraction occur as ≈≈ 200 390 285 364 410 cu 70 50 cu (at.%) 30 10 au 400 800 1000 t °c 1083 910 au,cu a uc uaucu₃ au₃cu 1063 90 fig. 35 one column fig. 35 binary phase diagram for au-cu (savitsky 1984; okamoto et al. 1987) showing the stability fields of phases with falling temperature. all these phases have been found in the hs concentrates from the skaergaard pge-au mineralisation. grey shading: temperature range for liquidus to solidus transition for (cu,au) alloy. solid lines: established boundaries of stability fields. dashed lines: approximate positions of the boundaries. boundaries at lower temperatures separate structural types of (cu,au) alloys stable at decreasing temperature (see okamoto et al. 1987 for details). table 22 percentage of precious metal mineral grains related to sulphides, silicates and feti oxides, and as free grains in the –125 mm fraction of hs concentrates sample no./ average centre/margin of intrusion mineralisation level number of grains sulphide % silicate and oxide % free % 18-958 centre au-rich au+1 33 31.2 5.3 63.5 24-1018 centre au-rich pd1/au 320 32.1 4.7 63.2 24-1022 au centre au-rich pd1 64 23.4 22.9 53.7 24-1022 pd centre pd-rich pd1 108 81.9 0.00 18.1 10-434 margin au-rich pd4 89 36.6 10.3 53.1 23a-798 margin au-rich pd4 506 10.8 26.8 62.4 18-1010 centre pd-rich pd5 180 77.1 3.9 18.9 24-1057 centre pd-rich pd5 373 72.2 0.7 27.1 23a-808 margin pd-rich pd5 116 56.3 10.2 33.5 10-445 margin pd-rich pd5 23 79.4 0.00 20.6 average centre au-rich pd1 to au+1 417 28.9 11.0 60.1 average margin au-rich pd4 595 23.7 18.6 57.8 average centre pd-rich pd5 561 77.1 1.5 21.4 average margin pd-rich pd5 139 67.9 5.1 27.1 sum of grains in sulphide, silicate and feti oxides and number of free grains = 100%. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 55 of 66 www.geusbul let in.org 0 0–5 5–10 10–15 15–20 20–25 25–30 30–35 35–40 40–45 45–50 50–55 55–60 60–65 65–70 70–75 75–80 80–85 85–90 90–95 95–100 2 4 6 8 10 12 14 16 18 % a 0 5 10 15 20 25 0–5 5–10 10–15 15–20 20–25 25–30 30–35 35–40 40–45 45–50 50–55 55–60 60–65 65–70 % b 0 5 10 15 20 25 30 35 40 45 0–5 5–10 10–15 15–20 20–25 25–30 30–35 35–40 40–45 45–50 50–55 55–60 60–65 % c fig. 36 one column tof n = 2829 average ecd = 23.6 ± 0.5 drill core 90-23a n = 533 average ecd = 19.3 ± 0.8 drill core 90-10 n = 161 average ecd = 22.5 ± 1.5 grain size (μm) fig. 36 variation in grain size (μm) of precious metal grains in a: central part of the mineralisation (tof: toe of forbindelsesgletscher), b: drill core 90-23a near the eastern margin and c: drill core 90-10 near the western margin. the central part of the mineralisation shows a relatively small increase in grains >30 µm. grains that are intergrown with sulphides (type i), or with rock-forming minerals of the gabbroic host (type ii) or as free grains (type iii; table 22; fig. 37). type i intergrowths and type iii free grains of precious metal account for 73% to almost all the precious metal phases in the studied samples (table 22; fig. 37). the proportion of free grains (type iii) reaches 63.5% in the au+1 mineralisation level in drill core 90-18. the highest proportions of precious metals attached to feti oxides and silicates are found in the au-rich mineralisation levels, represented by au concentrations in samples 24-1022 (23.4%) and 23a-798 (26.8%). in general, sulphide intergrowths dominate in samples from pd-rich mineralisation levels and vary from 56.3% to 79.4%, with an average of 77% in more central parts of the mineralisation and 68% toward the margins (table 22). the hrxrt investigations of godel et al. (2014) showed free grains of au-phases between rock-forming minerals in au-rich gabbro. gold tends to form free interstitial grains in correlation with elevated bulk-rock au, te and se (e.g. holwell et al. 2015; keays & tegner 2016), and their accumulation from residual melts and fluids (nielsen et al. 2015). in central areas the average is 60.1%, and 57.8% in the au-rich pd4a mineralisation levels near the margins (table 22). a l +sul +sil(ox) 24-1057 pd5 b 24-1018 pd1/au +sil(ox) +sul l c 24-1022 pd1 pd grains d 24-1022 pd1 au grains e 18-1010 pd5 f 18-958 pd1/au g 23a-808 pd5 h 23a-798 pd4 i 10-445 pd5 +sil(ox) l +sul fig. 37 1 column j 10-434 pd4 +sil(ox) +sul l i ii iii fig. 37 petrographic associations of precious metal mineral grains in individual samples (1 m intervals) from the lower pge-rich (pd5) mineralisation level and the au-rich upper mineralisation levels (see chapter 2) in –125 μm powders. samples are identified by the unique number of the drill core followed by the depth in the core from where the sample was collected. e.g. sample 24-1057 collects the interval between 1057 and 1058 m in drill core 90-24. the precious metal phases are as follows: i: attached to sulphides (+sul), ii: attached to silicates and feti oxides (+sil(ox)) or iii: as liberated grains (l). proportions of i to iii are normalised to 100% and include data from drill cores 90-10, 90-18, 90-23a and 90-24 (table 22). in the pge-rich layer (pd5) the precious metal minerals are associated with sulphides and occur as free phases with only small proportions associated with silicates and feti oxides. conversely, the precious metal grains in the au-rich mineralisation levels occur mostly as free grains but with high and variable proportions associated with either silicates and feti oxides or sulphides. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 56 of 66 www.geusbul let in.org pge-phases tend to relate to or be enclosed in sulphides while au phases tend to occur as free grains or are attached to silicates and feti oxides. this also applies to individual mineralisation levels and is highlighted by the lower pd-rich and upper au-rich fractions of sample 24-1022, representing the 1 m interval from 1022 m to 1023 m in the drill core (table 22, rows 5, 6; fig. 37c, d). based on the grain-size information in table 3, it is recommended that the mineralised gabbros are crushed to a grain size of c. 70 µm before floatation. additional magnetic separation could concentrate magnetite and ilmenite to recover sulphide droplets and precious metals trapped in feti oxides and potentially valuable biproducts such as ti, v and ga (rudashevsky et al. 2014, 2015). https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 57 of 66 www.geusbul let in.org 6 the mineralisation processes sulphides can be found throughout the gabbros of the skaergaard intrusion. in the lower part of the intrusion and prior to the formation of the pge-au mineralisation, the sulphides crystallised in interstitial blebs of trapped melt (e.g. wernette et al. 2020). the amounts of trapped melt and volatile-rich fluid can be estimated (e.g. tegner et al. 2009), but do not account for the proportions of the host rocks that crystallised from interstitial mush melts and whose residuals were subsequently recycled into the overlying bulk magma (nielsen et al. 2019b). precious metal bulk-rock anomalies in the upper levels of lzc and in mz, can be correlated to the occurrence of gabbroic and basaltic blocks from the roof zones and an associated increase in crystallisation of feti oxides (nielsen et al. 2015). increase in feti oxide crystallisation may have caused local sulphide saturation and increases in concentrations of precious metals (bird et al. 1991). irrespectively, the pd/pt ratios in gabbros just below the precious metal mineralisation has only increased from c. 3 to 6. a high d-value for pd compared to pt would have led to systematically decreasing pd/pt ratios in the remaining bulk liquid. instead, the increase in pd/ pt is due to limited solubility of pt compared to pd (nielsen et al. 2019b). below and prior to the formation of the pge-au mineralisation, sulphide saturation was reached in trapped liquids. in the upper mz and the mineralisation, sulphide saturation was reached in melts that crystallise and evolve to residuals that were recycled to overlying mush or bulk liquid. sulphide saturation in bulk liquid was reached during the formation of the lower uza (nielsen et al. 2015). major differences are apparent between the centre and the margins of the intrusion, despite very similar elemental distributions in mineralisation levels pd6 and pd5 (andersen et al. 1998; nielsen et al. 2005). at the margins, the entire precious metal mineralisation is hosted in pd6, pd5 and pd4 over a total width of c. 15 m and in two macrorhythmic layers. this compares to 60 m and five macrorhythmic layers in the central part of the mineralisation. total pge, calculated as ‘grade times width’ (g*w; m x ppm), is 20 at the margins compared to 35 in the centre (nielsen et al. 2019a). total au g*w is 4 near the margins and 12 in the centre (nielsen et al. 2019a). much less precious metal is found at the margins despite similar stratigraphic thicknesses of the gabbroic hosts. consequently, and during the formation and crystallisation of the macrorhythmic layers of the upper mz, a process concentrated the precious metals, especially au, progressively higher up in the stratigraphy and more centrally in the bowl-shaped, layered gabbros. the syn-crystallisation mobility of the precious metals in the macrorhythmic layers is key to the understanding of the processes that controlled the formation of the skaergaard pge-au mineralisation. here, we present a 10-step summary that integrates the mineralogical information presented here with the mineralisation model developed by nielsen (2016) and nielsen et al. (2015, 2019a) and the model for the evolution of the bulk liquid in nielsen et al. (2019b). we assume that cooling was most effective at the roof of the shallow crustal magma chamber, and that mushes formed under the roof descended to the floor in convective currents. the mushes crystallised in situ in the floor and recycled residual, evolved and buoyant melts to the bulk magma (nielsen et al. 2015, 2019a, b). the sequence of events that led to the observed mineralisation is therefore divided into two. firstly, processes that extracted all precious metals from remaining bulk liquid and concentrated them in the ‘mushy’ floor (steps 1–3; fig. 38a), and secondly the processes that led to the fully developed mineralisation in the floor mushes in the central parts of the intrusion (steps 4–10; fig. 38c). processes in the faster-cooled mush near the margins of the intrusion are subsequently discussed. steps 1 to 10 for the central parts of the intrusion are illustrated in fig. 38 and described as follows: steps 1–3: processes that concentrated precious metals in the floor mush: 1. once the entire magma chamber was established, the precious metals and volatiles accumulated in the remaining bulk magma like any other incompatible element. exceptions are precious metals and volatiles in trapped melts, and precious metals hosted in sulphides that formed as the result of the reaction between bulk liquid and roof blocks sinking into the magma chamber (nielsen et al. 2015). 2. the magma cooled from the roof, walls and floor. sulphide saturation was first reached in mush melt under the roof, down the walls and subsequently in the mush floor. at this stage, the bulk magma had feo* c. 22 wt%, pt 10 ppb, pd 63 ppb, au 42 ppb and cu 700 ppm. the mush liquid would have feo*c. 22 wt% (compatible element), pt 15 ppb, pd 94 ppm, au 63 ppm and cu 1000 ppm. 3. dense sulphides and immiscible fe-rich silicate melts that formed in the mush below the roof of the magma chamber were carried to the intermittent floor of the magma chamber in convective currents. in the floor, they mixed with residing mush. almost all precious metals ended up dissolved in the melts of the mushy floor. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 58 of 66 www.geusbul let in.org steps 4–10: sulphide saturation , dissolution of sulphides melt and redeposition of precious metals in floor mushes: 4. sulphide saturation was reached at t = 1030–1050°c, when a maximum of c. 65 vol.% of mush melt remained in macrorhythmic layer ml0. many tiny droplets of sulphide melt in suspension equilibrated with the remaining bulk liquid and preserved the pd/ pt ratio of the bulk melt (pd/pt = 6–10). only accumulated and protected sulphide and sulphide droplets that were enclosed in other minerals or otherwise protected from subsequent equilibration with interstitial melt preserved high pd/pt ratios. the pd/pt ratios are approximately five times that of the coexisting silicate liquid and reflect the high d-value of pd between sulphide and silicate liquids. 5. the mush melt reached the two-liquid field between si-rich and fe-rich melts at c. 1015°c, when a maximum of c. 40% melt remained in the mushes of the macrolayers, and buoyant, immiscible si-rich melt started to be recycled to the overlying bulk magma. 6. unprotected sulphide droplets redissolved in contact with the residing fe-rich melt, and redissolved cu, pt and au became available for upward migration in the remaining silicate melt and coexisting supercritical aqueous fluids. the dissolution process is similar to that of peregoedova et al. (2004). 7. within each macrorhythmic layer, decreasing temperature led to exsolution of precious metal phases from protected and unprotected sulphide droplets. 8. remaining mush melt enriched in volatiles, au and high field strength elements (hfse) migrated remaining bulk liquid host rocks layered series steps 4–10: accumulation of crystal mush in �oor mbs ubs mbs steps 1–3 mush 4 5 6–8 10 fractionation + sulphide droplets 2-liquids compaction i due to loss of granophyre compaction ii due to loss of hydrous and au-bearing residual melt mineralisation steps in macrorhythmic layers pxpl ag m el t establishment of macrorhythmic layers due to sinking and �oating of crystals mg lmush g mg lg tz tz pd 5 ml0 ml1.1 ml1.2 ml2.1 ml2.2 ml-1 0 50 60 30 leucogabbro au+1 pd2 pd3 pd4 pd4b a b c ml2 10 40 20 m ma�c gabbro gabbro pd6 pd1 pd2b pd3a pd5 pd1/au boundaries between macrorhythmic layers mineralisation levels in macrorhythmic layers (schematic) lg fig. 38 2 columns fig. 38 summary of the mineralisation processes in the skaergaard pge-au mineralisation, as developed in nielsen (2016) and nielsen et al. (2015, 2019a, b). a: trapping and accumulation of precious metals (steps 1–3) during which the precious metals are brought to the floor mush (based on fig. 26 in nielsen et al. 2015). blue: solidified gabbro. orange: mush zones with trapping. white: remaining bulk liquid. purple: palaeogene and precambrian host rocks. red arrows point to the zone of trapping of precious metals under the roof (steps 1–3) and the crystallisation and mineralisation processes in the floor mush (steps 4–10). b: correlation between mineralisation levels and the macrorhythmic layers. the bowl shape of the layers is omitted for clarity. the lower four macrorhythmic layers host two mineralisation levels each, including a lower level due to density accumulation and an upper level due to in situ sulphide saturation in residual fe-rich mush melt (based on fig. 8 in nielsen et al. 2015 and fig. 5 in nielsen et al. 2019a). c: upward accumulation, fractionation and transport of the precious metals in mush melts. steps 4–10 within ml0 macrorhythmic layer hosts the main pge-mineralisation level, pd5. sinking of dense phases and the rise of plagioclase in the proto-macrorhythmic layers allowed mush melt to evolve in between the two. sulphide saturation is reached due to crystallisation and fractionation of the mush melt (yellow dots with red rims) and is followed by silicate immiscibility (blue dots with black rims), and loss of si-rich conjugate melt, ponding of dense fe-rich silicate melts and subsequently the loss of the residual fe-rich silicate melt and free volatile phase. steps 4–10 are repeated in all mineralisation layers from ml0 to ml2.1 (based on fig. 22 in nielsen et al. 2015). mbs: marginal border series. ubs: upper border series. ml-1 to ml2.2: macrorhythmic layers. plag: plagioclase. px: pyroxene. mg: melagabbro. g: gabbro. lg: leucogabbro. tz: transitional zone. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 59 of 66 www.geusbul let in.org upwards in the overlying proto-macrorhythmic layer (nielsen et al. 2015). a small proportion of residual melt and fluid was trapped in the upper part of the macrorhythmic layers, as reflected in locally increasing au/pd ratios and in common hydrothermal alteration (nielsen et al. 2015). 9. processes in steps 4–8 are incrementally repeated in each of the macrorhythmic layers (ml1.1, ml1.2, ml2 and ml2.1; nielsen et al. 2019a). at each step, fluid content increases, and so do ratios of au/pd and pt/pd due to high d-values for pd. 10. residual fluid enriched in au, te and se (see section 5.4) rose to form the pge-poor uppermost mineralisation levels pd1/au and au+1. the paragenetic and petrographic relations are very different near the margins of the intrusion and reflect local geochemical environments (nielsen et al. 2019a). zvyagintsevite (pd3pb) dominates toward the eastern margin, and pd-sulphides dominate toward the western margin. at both margins, the relative proportions of these minerals decrease over a height of 15 m and up ml0. the plumbides and sulphides are followed first by arsenides, and subsequently by tellurides and (au,cu,ag) alloys. we note that the detailed description of parageneses at the eastern margin in holwell et al. (2016) does not represent the bulk of the mineralisation. droplets of cu-fe sulphide melt occur in feti oxides in mineralisation levels near the margins. the droplets testify to in situ sulphide saturation and processes similar to those in the centre. in the centre, the droplets are composed of bornite and chalcocite, but at the margins they are composed of bornite and chalcopyrite and represent a more s-rich paragenesis. examples of these droplets with pd sulphides and of compounds of ag, cu and au are shown in figs 23j, k, 24c, d, 27h. in samples from near the margins of the intrusion, the precious metal phases are not necessarily directly related to grains of cu-fe sulphide but are generally spatially related to h2o-bearing silicates and other lateto post-magmatic minerals including hornblende, biotite, actinolite, chlorite, talc, magnetite and quartz (figs 17l, 20l–n, 26h, 27o, p, 30b, c, e, f, 32e, i). grains of solid solutions of au, ag, cu and pd in these parageneses often have elongated tabular or pseudo-prismatic shapes. they fill the interstices between lateto post-magmatic silicates (figs 23l–n, 24k). the lastformed low-t precious metal phases are rare native gold and silver (figs 26j, 32e, g). at the eastern margin, towards the palaeogene lavas, the precious metal paragenesis comprises a polymineral association of pd-intermetallic compounds dominated by zvyagintsevite, with keithconnite, atokite and arsenides dominated by arsenopalladinite and vincentite and minor palladoarsenide, sperrylite and solid solutions of au, cu, and ag. the pd-intermetallic compounds are of very minor importance at the western margin, towards the precambrian basement where the pd-sulphides vysotskite and vasilite dominate, accompanied by the pd-arsenide vicentite, the pd intermetalloids keithconnite and (au,cu,ag,pt) alloy. these complex precious metal parageneses equilibrated with dense hydrothermal fluids trapped within the gabbros (figs 26g–j, n, o, 27 k–p, 30b–h, k, l). they recall low-t parageneses of low-sulphide ni-cu occurrences (spiridonov 2010b) and are most likely comparable to those of the monchegorsk magmatic layered complex and the monchetundra deposit, kola peninsula, russia (grokhovskaya et al. 2009). the low temperature of equilibration of the hydrous silicate paragenesis, the paragenetic differences and lateral variation in, for example, pge, au, ag, te and cu within a single stratigraphic layer of the host gabbros, suggest local controls by the coexisting fluids. the fluids were s-rich toward the precambrian basement to the west and s-poor toward the lavas at the eastern margin. the precious metal paragenesis crystallised and re-equilibrated in equilibrium with metal-rich fluids trapped in the crystallising gabbros. these fluids were out of equilibrium with the liquidus paragenesis of the gabbroic host. fluids as well as immiscible si-rich melts in paired melt inclusions were trapped in these gabbros (holness et al. 2011) and imply reduced permeability in gabbros that cooled close to the margins of the intrusion. the stratigraphic variations and levels of elemental concentrations in pd5 remain the same in drill cores across the intrusion (andersen et al. 1998). consequently, the processes that concentrated precious metal in the floor mush apply to all parts of the intrusion and steps 1–3 of the mineralisation processes therefore apply to the eastern as well as the western margins. however, due to reduced permeability and mobility of interstitial melts and aqueous fluids at the margins, the subsequent steps differ as outlined here: • the mush melt reached the two-liquid field between si-rich and fe-rich melts at c. 1015°c. fluids and silicate melts remained in the mush and equilibrium was maintained. • as crystallisation of the gabbroic host progressed, sulphide droplets evolved and exsolved precious metal at decreasing t (karup-møller et al. 2008). the parageneses equilibrated with trapped fluids and residual silicate melts resulting in hydrous silicate parageneses in matrix and veins. unprotected sulphides reacted and equilibrated with hydrous residual melts and trapped fluid. • zvyagintsevite (pd3pb) dominated the paragenesis at the eastern margin, and pd-sulphides dominated at the western margin. the paragenetic differences between the two margins reflect local geochemical environments https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 60 of 66 www.geusbul let in.org and seemingly also compositional differences between meteoric fluid from the precambrian basement versus that of the palaeogene lava successions. thus, precious metal paragenesis formed near the margins of the intrusion cannot represent the paragenesis of the bulk of the skaergaard pge-au mineralisation. • the proportions of zvyagintsevite at the eastern margin and of pd-sulphides at the western margin decrease upwards in pd5 and give way, first to pd-arsenides and subsequently to (au,cu,ag) alloys typical of parageneses in equilibrium with late magmatic hydrothermal fluids. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 61 of 66 www.geusbul let in.org 7 conclusions the contrast between precious metal parageneses and grain morphologies in the lower pge-rich and the au-rich upper mineralisation levels confirms the orthomagmatic origin of the skaergaard pge-au mineralisation. pd was concentrated in droplets of an immiscible cu and fe sulphide melt, aided by subsequent resorption. whereas the late accumulation of au was the result of upwards hydrothermal transport. the mineralisation processes described here span a wide temperature interval from c. 1050°c to temperatures below the solidus. the observed distribution of precious metals in the skaergaard pge-au mineralisation is the result of the interplay between (1) the bulk composition of the magma, (2) the solubilities of the precious metals, copper and sulphur, (3) the density control on accumulation and mobility of melts and solids and (4) contrasts in cooling conditions in the magma chamber. fundamental to the mineralogy, morphology and petrographic relations of the precious metal parageneses, is the timing of sulphide saturation and the mobility of residual melts and fluids. immiscible droplets of cu and fe sulphide formed throughout the residual mushes of the macrorhythmic layers at temperatures of 1050 to 1030°c. some droplets accumulated and were protected in the lower mineralisation levels, and the remainder disseminated in the upper parts of individual macrorhythmic layers. subsequent silicate immiscibility at c. 1015°c led to resorption and syn-magmatic rather than post-solidification loss of s. the sulphide droplets exsolved precious metals in droplets of melts that crystallised as, for example, ‘high skaergaardite’. cooling also led to diffusion of precious metals from sulphides to the remaining mush melts and fluid, which were then made available for upward migration. all sulphides and precious metal phases equilibrated with the remaining mush melts and fluids as temperatures dropped following known phase relations. these processes led to the stratification of the mineralisation and the fractionation of upward-decreasing pd/pt ratios and increasing au/pd ratios, and the total dominance of skaergaardite (pdcu) with upward-increasing au in the bulk of the skaergaard pge-au mineralisation. at the margins of the intrusion, faster cooling reduced permeability, hence the mobility of residual melts and fluids was hampered. this resulted in the absence of resorption of sulphide droplets and the preservation of a bornite-chalcopyrite sulphide paragenesis. only residual fluids containing au and elements carried in high-t magmatic systems escaped from the lowermost macrorhythmic layer. subsequent recrystallisation and equilibration to hydrous, low-t parageneses with sulphides, arsenide and intermetallic compounds was controlled by local geochemical constraints, comparable to late magmatic parageneses known from pge and au deposits elsewhere. the skaergaard pge-au mineralisation is seemingly unique in its extensive preservation of sulphide droplets the subsequently crystallised precious metal minerals and the preservation of exsolved droplets of precious metals. as such, the mineralisation provides insights into the early stages of precious metal mineralisation in layered intrusions that are often obliterated by subsequent hydrothermal reactions at lower temperatures and remobilisation of the metals. our detailed record of the mineralogy and petrographic relations across the skaergaard intrusion testify to the significant complexity in metal distributions, mineralogy and genesis of skaergaard-type pge-au mineralisations. almost all models proposed and demonstrated for the formation of precious metal mineralisations in layered gabbro complexes apply to one or another part or formation stage of the skaergaard pge-au mineralisation. no classic orthomagmatic model, nor a model advocating fluid redistribution of precious metals from already solidified parts of the intrusion can account for our observations, but a combination of the two models can. the concentration of the precious metals is the result of syn-magmatic sulphide saturation below the roof of the intrusion, followed by repeated sulphide saturation in floor mushes, re-equilibration and redistribution by mush melts and fluids in the accumulating gabbros. additional information funding statement these investigations were co-financed by the geological survey of denmark and greenland (geus) and cnt instruments llc. author contributions nsr and vnr: laboratory analyses, preparation of reports on which this study is based, writing of the manuscript. tfdn: project design, selection and collection of samples, development of petrogenetic models, editing of the manuscript and main author of the introduction, discussion and conclusions. competing interests the authors declare no competing interests besides that nikolay s. rudashevsky and vladimir n rudashevsky are affiliated with the commercial company, cnt instruments llc. acknowledgements these investigations could not have been carried out without the continued support by late dr. leif thorning, head of the department for petrology and economic geology, geus. guest editor rune b. larsen and geus bulletin editor in-chief catherine jex, are thanked for their most helpful comments and advice. reviews by frederica zaccarini and eduardo mansur were of  great help during revision of the manuscript. artwork was produced with valuable help from susanne rømer and the drawing office of geus. https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org rudashevsky et al. 2023: geus bulletin 54. 8306. 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https://doi.org/10.1007/bf02645723 https://doi.org/10.34194/geusb.v26.4754 https://doi.org/10.1093/petrology/egp020 https://doi.org/10.2138/am.2013.4044 https://doi.org/10.1007/s00126-019-00924-1 https://doi.org/10.1016/j.epsl.2012.08.043 https://doi.org/10.1016/j.oregeorev.1042573 https://doi.org/10.1016/j.oregeorev.1042573 https://doi.org/10.1007/s00710-004-0049-1 https://doi.org/10.1016/j.gca.2013.02.026 https://doi.org/10.1016/j.gca.2013.02.026 rudashevsky et al. 2023: geus bulletin 54. 8306. https://doi.org/10.34194/geusb.v54.8306 66 of 66 www.geusbul let in.org appendix 1 formulae for common host-rock minerals and base metal sulphides rock-forming minerals olivine: (mg,fe)2sio4  forsterite: mg2sio4  fayalite: fe2sio4 pyroxene group  orthopyroxene: (mg,fe)sio3  clinopyroxene: ca(mg,fe)si2o6 plagioclase: (na,ca)[(si,al)alsi2]o8  anorthite: caal2si2o8  albite: naalsi3o8 feti oxides  titanomagnetite: fe2+(fe3+,ti)2o4  ilmenite: fetio3  magnetite: fe3o4 apatite ca5(po4)3(cl/f/oh) biotite: k(mg,fe)3alsi3o10(oh)2 (simplified) amphibole group   hornblende: (ca,na)2(mg,fe,al)5(al,si)8o22 (oh)2  (simplified)  actinolite: ca2(mg,fe)si8o22(oh)2  tremolite: ca2mg5(si8o22)(oh)2  cummingtonite: mg7(si8o22)(oh)2  pargasite: naca2(mg4al)(si6al2)o22(oh)2 talc: mg3si4o10(oh)2 chlorite: (mg,fe)3(si,al)4o10(oh)2·(mg,fe)3(oh)6 serpentine: (mg,fe,ni, mn,al,zn)3[si2o5](oh)4 ferrosaponite: ca0.3(fe2+,mg,fe3+)3((si,al)4o10)(oh)2 · 4h2o calcite: caco3 quartz: sio2 rare accessory minerals baddeleyite: zro2 zircon: zrsio4 thorianite: tho2 monazite: ree(po4) uraninite: uo2 thiospinel: variety carrollite cu(co,fe,ni)2s4 base metal sulphides bornite: cu5fes4 chalcocite: cu2s digenite: cu9s5 chalcopyrite: cufes2 pyrrhotite: fe1-xs galena: pbs sphalerite: zns pentlandite: (fe, ni)9s8 cobaltian pentlandite: (co,ni,fe)9s8 https://doi.org/10.34194/geusb.v54.8306 http://www.geusbulletin.org the pge-au mineralisation of the skaergaard intrusion: precious metal minerals, petrography and ore 1 introduction 2 the skaergaard intrusion 2.1 the skaergaard pge-au mineralisation 2.2 the bowl-shape of the triple group 3 samples and methods 4 mineralogy, composition and petrographic relationships 4.1 the gabbroic host 4.2 the ore minerals 4.2.1 the sulphides 4.2.2 precious metal minerals 4.2.3 intermetallic compounds and alloys of cu and pd 4.3 relationships between host rocks, sulphides and precious metal phases 4.3.1 droplets and globules 4.3.2 cu-pd minerals 4.3.2.1 relationships between cu-pd minerals and cu-fe sulphides 4.3.2.2 intergrowths of more cu-pd minerals 4.3.2.3 relationship between cu-pd minerals and other precious metal phases 4.3.2.4 pt minerals 4.3.3 intermetallic compounds and alloys of au, cu, and ag 4.3.3.1 relationships of intermetallic compounds and (au,cu,ag) alloys to cu-fe sulphides 4.3.3.2 grain intergrowths of more au-cu minerals 4.3.3.3 relationships of intermetallic compounds and (au,cu,ag) alloys with pges 4.3.3.4 relationships between au-cu compounds and rock-forming minerals 4.3.4 pge sulphides and s-bearing ag minerals 4.3.5 intermetallic compounds of pd, cu and ag with pb, te, sn, sb and bi 4.3.6 arsenides of pd, pt, and ni 5 discussion: constraints for the mineralisation model 5.1 concentrations of pge and au in the magma at sulphide saturation 5.2 constraints on the timing and temperature of sulphide saturation 5.3 petrographic evidence for an initial high-t origin of the mineralisation 5.4 the role of coexisting fluids in the skaergaard pge-au mineralisation 5.4.1 evidence for coexisting magmatic fluid 5.4.2 nature of the coexisting fluid phase 5.4.3 extent of equilibration with aqueous fluids 5.5 low-t re-equilibration 5.6 petrographic information relevant to beneficiation of the ore 6 the mineralisation processes 7 conclusions additional information acknowledgements references appendix 1 formulae for common host rock minerals and base metal sulphides rock-forming minerals rare accessory minerals base metal sulphides figures fig. 1 overview of the skaergaard intrusion. a: simplified geological map slightly modified from nie fig. 2 the leucogabbro layers (l1, l2 and l3) of the triple group on wagertoppen. a: l1 to l3 on the fig. 3 summary of the skaergaard mineralisation. from left to right: macrorhythmic layers (ml-1 to m fig. 4 correlation between elemental concentrations and lithology of host gabbro in the skaergaard p fig. 5 visualisation of the 3d structure of the precious metal-rich part of the mineralisation (not fig. 6 backscattered electron (bse) images of the petrographic relations between rock-forming silica fig. 7 bse images of sulphides. panels a-e are polished thin sections of gabbroic host. panels f-n a fig. 8 bse images of rare minerals in polished monolayer thin sections. table 3 provides formulae fo fig. 9 bse images of sulphides from the au-rich pd1/au horizon in the central part of the intrusion. fig. 10 contrasts between precious metal parageneses across the intrusion in the upper au-rich miner fig. 11 intermetallic compounds and alloys ranging from cu to pd. based on emp analyses of a total o fig. 12 factor analysis on the skaergaardite (pdcu) composition from sample 90-24-1057 in level pd5. fig. 13 histograms illustrating frequency of substitutions in skaergaardite, based on a total of 148 fig. 14 bse images of silicate globules in ilmenite and examples of hydrous silicates coexisting wit fig. 15 re-compiled high-resolution x-ray computed tomography images (hrxct) after godel et al. (201 fig. 16 bse images showing the petrographic relations between skaergaardite and cu-fe sulphides in p fig. 17 bse images showing the petrographic relationships between nielsenite (cu3pd), (pd,cu) alloys fig. 18 bse images of free grains of cu and pd minerals in polished monolayer thin sections of hs co fig. 19 bse images showing the relationship between cu and pd minerals and other precious metal mine fig. 20 bse images of pt minerals in polished monolayer thin sections of hs concentrates. table 3 pr fig. 21 compositions and proportions of au-cu minerals and phases. the histogram shows the wide comp fig. 22 au-rich parageneses in a: mineralisation level pd1/au in all drill core samples from the cen fig. 23 bse images showing the petrographic relations between (au,cu) phases and rock-forming minera fig. 24 bse images of free grains of au and cu minerals in monolayer polished thin sections of hs co fig. 25 bse images of zoned grains of au minerals, cu minerals and exsolution textures in polished t fig. 26 bse images of the petrographic relations between au minerals and cu minerals with pges in po fig. 27 relationship between au and cu minerals and rock-forming minerals of the gabbroic host in po fig. 28 bse images of vysotskite and vasilite in polished monolayer thin sections of hs concentrates fig. 29 bse images of rare pge sulphides in polished monolayer thin sections of hs concentrates. all fig. 30 bse images of pd, cu and ag intermetallic compounds with pb, te, sn, sb and bi in polished m fig. 31 bse images of rare intermetallic compounds of precious metals in polished monolayer thin sec fig. 32 bse images of pge arsenides in polished monolayer thin sections of hs concentrates. as for t fig. 33 phase diagrams showing compositions (wt%) of a: primary cu, pd and au and b: primary alloys fig. 34 binary phase diagrams for cu and pd after savitsky (1984) and baker et al. (1992). a: stabil fig. 35 binary phase diagram for au-cu (savitsky 1984; okamoto et al. 1987) showing the stability fi fig. 36 variation in grain size (mm) of precious metal grains in a: central part of the mineralisati fig. 37 petrographic associations of precious metal mineral grains in individual samples (1 m interv fig. 38 summary of the mineralisation processes in the skaergaard pge-au mineralisation, as develope tables table 1 specifications of the studied samples table 2 composition of rare sulphides, sulpho-tellurides, tellurides and arsenides table 3 precious metal minerals of the skaergaard pge-au mineralisation table 4 average compositions of cu and pd minerals table 5 correlation coefficients for elements contained in 1482 grains of skaergaardite table 6 composition of varieties of skaergaardite and nielsenitea table 7 average composition of skaergaardite in mineralisation levels pd5-pd4 and pd2-pd1 table 8 average composition of pt alloys table 9 average compositions of au, cu, pd and ag minerals table 10 compositional range of au-cu mineralsa table 11 average compositions of au and cu minerals per mineralisation level table 12 compositional zonation in au-cu mineral grains table 13. composition of au-cu minerals and their exsolution phases table 14 compositional types of vysotskite and vasilite in the skaergaard pge-au mineralisation table 15. compositions of rare precious metal sulphides table 16 compositional types of zvyagintsevite in the skaergaard au-pge mineralisation table 17 compositional types of keithconnite and kotulskite table 18 compositional types of atokite and cabriite table 19 compositions of naldrettite, sopcheite, froodite, merenskyite, hessite and unnamed pd2te, p table 20 compositional types of arsenopalladinite and palladoarsenite table 21. compositional types of vincentite/guanglinite and compositions of isomertieite, majakite a table 22. percentage of precious metal mineral grains related to sulphides, silicates and feti oxide larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 1 of 121 *correspondence: lml@geus.dk received: 21 sept 2021 accepted: 19 apr 2022 published: 10 oct 2022 keywords: palaeogene, flood basalts, vaigat formation, svartenhuk formation, naqerloq formation abbreviations geus: geological survey of denmark and greenland gps: global positioning system icp-ms: inductively coupled plasma mass spectrometer loc.: locality odp: ocean drilling program ree: rare-earth element(s) tas: total-alkalis-silica utm: universal transverse mercator xrf: x-ray fluorescence geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed jornal published by the geological survey of denmark and greenland (geus) this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the orriginal work. author(s) retain copyright. edited by: adam a garde (geus, denmark) reviewed by: david peate (university of iowa, usa), j. richard wilson (aarhus university, denmark) funding: see page 113 competing interests: none declared. additional files: see page 113 abstract the palaeogene volcanic succession in the northern part of the nuussuaq basin in west greenland comprises three formations: the vaigat and svartenhuk formations of paleocene age (61–58 ma) and the naqerloq formation of eocene age (57–54 ma). in this study, we formalise and describe the volcanic stratigraphy on svartenhuk halvø and the areas with lavas that flowed across the basin boundary onto the adjoining basement areas in the north and east. the vaigat formation comprises three members. the kakilisaat and nerutusoq members are of minor volume and consist of, respectively, crustally contaminated basalts and chemically enriched basalts with relatively high contents of incompatible trace elements. they are overlain by the voluminous nunavik member of tholeiitic picrites (mgo ≥12 wt%) and subordinate magnesian basalts. the oldest volcanic deposits are commonly foreset-bedded hyaloclastites, and the overlying subaerial lavas are mainly thin, grey, crumbling flows. eruption sites were mainly within the basin, with depocentres in the south and hyaloclastite and lava transport directions towards the north. thicknesses vary from up to at least 2000 m in the south to ≥380 m in the northernmost exposures close to 72°n. the svartenhuk formation comprises four members. the lowest, kuugaartorfik member, is up to 100 m thick and consists partly of quartzofeldspathic and partly volcanogenic sediments; it is restricted to northern svartenhuk halvø and the innerit peninsula. the overlying volcanic tunuarsuk, nuuit and skalø members are voluminous and widespread, with a combined thickness of up to 1800 m. they consist of tholeiitic basalts with similar chemical compositions but with correlatable stratigraphic variation patterns. the tunuarsuk member consists of interspersed flow groups of thin, grey flows and massive, brown flows; the nuuit member comprises mainly massive brown flows, and the skalø member is dominated by light grey flows. the svartenhuk formation oversteps the vaigat formation on the basement in the north and east. in these distal areas the tunuarsuk and nuuit members constitute the major volumes, and preserved thicknesses are up to 1400 m. in northern and eastern svartenhuk halvø and also farther to the north and east, foreset-bedded hyaloclastites indicate transport directions towards the north and possibly east from eruption sites within the basin. the naqerloq formation comprises one member, the arfertuarsuk member, consisting of flows of brown basalt with relatively enriched chemistry and a single trachyte flow. the member is only found in western svartenhuk halvø and on skalø, where it conformably overlies the older lavas with up to 350 m thickness preserved after erosion. dykes of all three formations are present. the distribution of dykes of the naqerloq formation suggests that this originally extended much farther east. picrites and basalts of the vaigat and svartenhuk formations are geochemically related; the picritic lavas represent erupted primitive magmas, whereas the basaltic lavas represent fractionated melts formed in deep magma chambers. the melts formed from a geochemically depleted but heterogeneous mantle; in addition melts from enriched sources were occasionally incorporated. the enriched basalts of the naqerloq formation arose from another mantle source. low contents of v, cu and ni in some crustally contaminated lavas indicate that accumulation of these elements may be present at depth. lithostratigraphy, geology and geochemistry of the tertiary volcanic rocks on svartenhuk halvø and adjoining areas, west greenland jørgen gutzon larsen1 and lotte melchior larsen 2* 1 fasanvænget 220, kokkedal, denmark. 2 geological survey of denmark and greenland (geus), copenhagen, denmark. monograph https://orcid.org/0000-0002-9344-4166 www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 2 of 121 ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ 200 km greenland ice sheet nuussuaq basin baffin bay davis strait hecla high sisimiut basin labrador sea 60°w 50°w 60°w 50°w70°w 74°n 70°n 66°n 66°n 74°n 70°n cape dyer delta-1 alpha-1 fig. 2 bb8-d12-d13 u ng av a fa ul t s ys te m disko nuussuaq svartenhuk halvø m elville bay basin ue gv01_08_014_01_lml_01.eps greenland mesozoic basin eocene oceanic crust paleocene oceanic crust palaeogene volcanic rocks shallow or exposed continental basement exploration wells faults, with downthrow directions extinct spreading centre transform fault dredge stations upernavik escarpmentue ■ unknown crustal type fig. 1 simplified geological map showing the regional and tectonic setting of the nuussuaq basin in west greenland. modified from fig. 1 in dam et al. (2009) with results from funck et al. (2012), oakey & chalmers (2012) and gregersen et al. (2019). dredge stations bb8-d12 and d13 on the upernavik escarpment from polteau & planke (2008). http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 3 of 121 introduction the nuussuaq basin is one of a series of linked basins that extend along the whole western continental margin of greenland with a basin fill comprising mesozoic to tertiary sediments and tertiary volcanic rocks (chalmers & pulvertaft 2001; fig. 1). the basins are mainly situated in the offshore areas, but between 69° and 73°n the rocks of the nuussuaq basin are exposed on disko (qeqertarsuaq), nuussuaq, ubekendt ejland (illorsuit) and svartenhuk halvø (nunavik or sigguup nunaa; halvø is peninsula in danish; fig. 2). the rocks have been studied since the mid-1850s because of occurrences of well-preserved fossils and coal in the sediments and of native iron in the volcanic rocks. in later years, hydrocarbons were discovered, mainly in the volcanic rocks. moreover, the structures, lithologies and compositions of the exposed parts of the basin have been studied in order to provide clues and correlations to the geology of the large offshore areas where drilling has been carried out. the volcanic rocks crop out over c. 20 000 km2 within a basinal area of c. 120 km × 400 km (c. 50 000 km2); including the sea-covered areas the volcanic rocks extend over a total area of c. 200 km × 550 km (c. 110 000 km2; chalmers et al. 1999; oakey & chalmers 2012). offshore, they extend into the melville bay basin in the north and the sisimiut basin in the south (fig. 1). the volcanic rocks in the nuussuaq basin were erupted at highly variable rates into a tectonically very active environment, and there was a complex interplay between the volcanic, sedimentary and tectonic evolution of the nn nuussuaq 71°n 72°n 72°n 51°w53°w 71°n 70°n 69°n 55°w 51°w svartenhuk halvø ubekendt ejland greenland ice sheet uummannaq innerit hareøen iti lli fa ult vaigat disko disko bugt ilulissat qeqertarsuaq aasiaat cbf cbf 50 km greenland neogene sediment cover o�shore sarqâta qáqâ central complex, ubekendt ejland vaigat formation undi�erentiated basalts o�shore naqerloq formation svartenhuk formation maligât formation maastrichtian– paleocene sediments albian–campanian sediments extensional fault precambrian basement fault with lateral or alternating displacements gb01_02_019_06_lml.eps fig. 2 simplified geological map of the nuussuaq basin. light colours: sea-covered areas. cbf: cretaceous boundary fault system. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 4 of 121 basin, which resulted in a complex architecture of the volcanic succession. moreover, the range of magma compositions was large. west greenland contains an unusually high proportion of primitive, mg-rich picritic rocks that represent almost unmodified mantle melts (drever 1953, 1956; clarke 1970; clarke & upton 1971; clarke & pedersen 1976; holm et al. 1992, 1993; larsen & pedersen 2000, 2009). however, repeated episodes of assimilation of crustal material, including organic-rich sediments, led to formation of units of silica-enriched volcanic rocks. on disko and nuussuaq these crustally contaminated rocks include native-ironand graphite-bearing basalts, andesites, dacites and rhyolites (treated in detail by pedersen et al. 2017, 2018). the volcanic succession is divided into two major parts. the lower part is dominated by picrites and has been formalised as the vaigat formation, which is present from disko and northwards to svartenhuk halvø and innerit (figs 2, 3; hald & pedersen 1975; clarke & pedersen 1976). the upper part is dominated by basalts and has been referred to different formations and members in different areas; according ? ? ? nûluk mb ordlingassoq mb b e l o w e x p o s u r e l e v e l ordlingassoq mb qeqertalik mb tuperssuartâta kûa mb 61.08 ± 0.56* 60.2 ± 0.5 niaqussat mb dykes 57.49 ± 1.40 58.34 ± 0.40 dykes 57.49 ± 1.40 58.34 ± 0.40 nordfjord mb rinks dal mb 61.2 ± 0.4 ordlingassoq mb 61.3 ± 0.5 tunoqqu & kûgá. mbs naujánguit mb 61.2 ± 0.5 anaanaa mb 62 (c27n) 58.66 ± 0.34 54.03 ± 0.33 48.02 ± 2.76 56.15 ± 0.41 58.31 ± 0.30 54.86 ± 0.32 alkaline sill 54.52 ± 0.67 dykes 54.3 ± 0.3 55.5 ± 0.4 56.8 ± 0.8 alkaline sill 54.52 ± 0.67 dykes 54.3 ± 0.3 55.5 ± 0.4 56.8 ± 0.8 gabbro sill 58.96 ± 0.51 gabbro sill 58.96 ± 0.51 avatarpaat plug 27.8 ± 0.6 h ar eø en fo rm at io n n aq er lo q fo rm at io n sv ar te nh uk fo rm at io n m al ig ât fo rm at io n va ig at fo rm at io n er qu a fm 39 –3 8 56 –5 4 60 –5 8 61 –6 0 eo ce ne pa le oc en e 62 –6 1 54 –5 3 talerua mb 38.74 ± 0.23 aumarûtigssâ mb dykes dykes 34.5 ± 0.2 ubekendt ejlanddisko nuussuaq ma hareøen and west nuussuaq west of itilli fault east of itilli fault svartenhuk halvø erqua fm 53.47 ± 0.52 55.91 ± 0.60 54.86 ± 0.44 54.40 ± 0.81* 55.20 ± 0.79* 55.31± 1.09* skalø mb 57.98 ± 0.59 nuuit mb 58.05 ± 0.59 tunuarsuk mb 59.41 ± 0.61 59.78 ± 0.41* 60.31 ± 1.39 kuugaar. mb 55.94 ± 0.20 u. m. l. u. m. l. s.q. gv02_02_047_01_lml 57.25 ± 0.95 56.99 ± 0.49 55.21 ± 0.30 59.50 ± 1.71 59.97 ± 0.89 kanísut mb nûk takisôq mb ifsorisok mb c 27 n c 26 r c 27 r c 26 r c 26 r c 24 r c 24 r c 18 n c 17 r nerutusoq mb kakilisaat mb nunavik mb arfertuarsuk mb fig. 3 stratigraphic scheme and 40ar–39ar radiometric ages (in ma) for the volcanic rocks in the nuussuaq basin, modified from larsen et al. (2016). radiometric ages with one digit after the decimal point are from storey et al. (1998) and larsen et al. (2009), ages with two digits after the decimal point (and no asterisk) are from larsen et al. (2016) and ages marked with an asterisk are from chauvet et al. (2019). the age for the anaanaa member is not radiometric but based on its normally magnetised character. note that the vertical ‘age scale’ is not equidistant. geological units are based on hald & pedersen (1975) for disko and nuussuaq, hald (1976) for hareøen and west nuussuaq, larsen (1977) for ubekendt ejland, and this study for svartenhuk halvø. s.q.: sarqâta qáqâ central complex; kuugaar. mb: kuugaartorfik member; l., m. and u.: lower, middle and upper nûk takisôq member, respectively. sediments with a quartzofeldspathic component are yellow and purely volcaniclastic sediments are brown. wavy lines indicate unconformities. narrow black and white columns on the right side of some lithological columns are normal and reversed palaeomagnetic directions, respectively, with magnetochrons indicated; data from riisager & abrahamsen (1999), riisager et al. (1999, 2003), schmidt et al. (2005), and p. riisager (unpublished data, 2006). http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 5 of 121 to the most recent dating and correlation by larsen et al. (2016), the maligât formation is present on disko and on nuussuaq east of the itilli fault, whereas the svartenhuk formation is present on nuussuaq west of the itilli fault, on ubekendt ejland, svartenhuk halvø and the areas farther north and east of this (figs 2, 3). in this study, a new basalt formation of early eocene age, the naqerloq formation, is established, which is present from hareøen in the south through western nuussuaq and ubekendt ejland to svartenhuk halvø in the north. the volcanic stratigraphy of hareøen and westernmost nuussuaq was described by hald (1976), that of ubekendt ejland by larsen (1977), and that of disko and nuussuaq by pedersen et al. (2017, 2018). in this study, we formalise and describe the volcanic stratigraphy of the northern part of the nuussuaq basin on svartenhuk halvø, including the lavas that flowed across the cretaceous boundary fault system to the adjoining areas in the north, north-east and east. the major part of the work was done south of 72°n on svartenhuk halvø and the southern part of the innerit peninsula in connection with mapping for the two 1:100 000 scale geological maps that cover these areas: 71 v.1 syd igdlorssuit and 71 v.1 nord svartenhuk (fig. 4). basalt flows occur as far north as 72°43′n and as far east as 51°54′w (fig. 5), but north of 72°n and east of 53°40′w, the data coverage is of reconnaissance character only. the extent of the volcanic rocks shown on the published 1:100 000 geological maps 71 v.2 nord nûgâtsiaq/nuugaatsiaq and 72 v.2 syd pangnertôq/ pannertooq (fig. 4) is mainly based on photo interpretation (henderson & pulvertaft 1987; guarnieri et al. 2022a, b). the area north of 72°n and west of 54°w is at present only covered by the 1:500 000 geological map sheet 4 upernavik isfjord (escher 1985). previous investigations of the svartenhuk halvø area the earliest scientific descriptions of the volcanic rocks in west greenland were published by giesecke (1823; 1910) and rink (1852) who used the terms “floetz-trap” or “trap-formation” and distinguished between “traptuff” (hyaloclastite), “basalt-tuff” (pillow breccia), “amygdaloid basalt” and “columnar basalt”. the first geological map was produced by steenstrup (1883) and shows the distribution of basement, sediments and volcanic rocks comprising “traptuf (palagonit?)” and “trap” between 69°10′n and 72°35′n, with the major parts of the inland areas still unmapped. in the first half of the 20th century, the general geology and structure of the disko–nuussuaq–svartenhuk region was described by koch (1929). nieland (1931) described the mineralogy of an anorthoclase-rich trachyte flow from arfertuarsuk and an olivine basalt dyke from south-east svartenhuk halvø and presented probably the first chemical analyses from the area. rosenkrantz et al. (1942) presented a reconnaissance geological description of svartenhuk halvø. a more detailed study of the volcanic rocks of svartenhuk halvø was published by noe-nygaard (1942), distinguishing for the first time older picrites (“oceanites”) from younger plagioclase-phyric basalts. noe-nygaard (1942, p. 67–68) estimated the thickness of the picrite succession exposed along the south coast of svartenhuk halvø to more than 10 km, based on the explicit assumption that there are no faults (no faults were observed due to foggy weather on the boat trip along the coast). major faulting was indeed demonstrated by later workers (rosenkrantz & pulvertaft 1969; münther 1973; larsen 1983; larsen & grocott 1991). after the second world war, new expeditions to the area led to increased understanding and detailed knowledge of both the nuussuaq basin and the volcanic rocks, see e.g. pulvertaft & clarke (1966) and münther (1973). early reviews on the stratigraphy, tectonics and structure of the basin were given by rosenkrantz & pulvertaft (1969), henderson (1973) and henderson et al. (1981). drever & game (1948) and drever & johnston (1957), working on the picrites on ubekendt ejland, were the first to point out that the picrites must represent highly magnesian and very hot magmas. this view was supported by clarke (1968, 1970), clarke & upton 71°n 72°n 51°w53°w 50 km gv01_02_181_lml pangnertôq/ pannertooq 72 v.2 syd igdlorssuit 71 v.1 syd mârmorilik/maarmorilik 71 v.2 syd svartenhuk 71 v.1 nord prøven 72 v.1 syd upernavik 72 v.1 nord nûgâtsiaq/ nuugaatsiaq 71 v.2 nord west greenland basalt group precambrian basement map sheets 1:100 000 nuussuaq group 72°n fig. 4 index map of geological maps covering svartenhuk halvø and the surrounding areas. maps with names given in both old and new greenlandic spelling were published 1970–1991 and in revised form in 2022 (guarnieri et al. 2022 a,b,c). the prøven and upernavik maps are in compilation. the igdlorssuit and svartenhuk maps (larsen 1983; larsen & grocott 1991) are provided in supplementary files s1 and s2. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 6 of 121 (1971) and o’nions & clarke (1972) working on baffin island and svartenhuk halvø. a comprehensive review of the volcanic rocks of west greenland was published by clarke & pedersen (1976). this review is still essentially correct although many details have been added since then and, in particular, the stratigraphy and geochemistry of the volcanic succession are now much better documented and understood. the two geological maps at scale 1:100 000 covering svartenhuk halvø were published by larsen (1983) and larsen & grocott (1991). the geological basis for the maps was described in a comprehensive unpublished report by larsen (1981a), and the present volume includes much information contained in that report without further reference to it. an important paper, to which this study may be seen as complementary, is that of larsen & pulvertaft (2000). for the first time, the complicated structure of svartenhuk halvø is here described and interpreted in detail, with faulting and tilting of the sedimentary-volcanic succession. the paper also gives short overviews of the lithology of both sediments and volcanics based on larsen (1981a) and later field work. consequently, here we only treat the structure of svartenhuk halvø briefly and repeatedly refer to larsen & pulvertaft (2000). in recent years, various aspects of the geology of svartenhuk halvø have been investigated. gill et al. (1992) and holm et al. (1992, 1993) presented petrological, geochemical and isotopic studies; geoffroy et al. (2001), abdelmalak et al. (2012) and chauvet et al. (2019) investigated the structure of the faulted areas; riisager et al. (2003) made palaeomagnetic profiles through the vaigat formation; dam et al. (2009) established the formal lithostratigraphy of the cretaceous– tertiary sediments; larsen et al. (2016) presented 40ar/39ar age determinations, and agranier et al. (2019) presented geochemical data for a large sample set with a petrological interpretation of the magma genesis for the area. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 7 of 121 geological setting general geology the volcanic succession on svartenhuk halvø and adjoining areas belongs to the west greenland basalt group (hald & pedersen 1975). this group extends onshore over a large area from disko (69°15′n) in the south to inland areas east of upernavik (72°43′n) in the north, a distance of nearly 400 km. it extends from the coastal areas and up to 150 km inland where basalt remnants cap the precambrian basement at altitudes up to nearly 2000 m a.s.l. the nuussuaq basin is bounded towards the precambrian basement areas in the east by a system of large, deep faults, which run irregularly nnw–sse and step repeatedly to the west (fig. 2). these are referred to as the cretaceous boundary fault system (larsen & pulvertaft 2000). this fault system runs through svartenhuk halvø and divides it into two parts: the cretaceous–tertiary basin in the west and the precambrian basement covered by basalts in the north and east (figs 2, 5, 6). the precambrian basement in the northern area considered here consists of archaean and proterozoic gneisses overlain by proterozoic metasediments and metavolcanics of the karrat group (e.g. grocott & pulvertaft 1990; thrane 2021; guarnieri et al. 2022c). in the north, these rocks are cut by a large intrusion, the prøven igneous complex, also of proterozoic age (e.g. escher 1985; thrane et al. 2005). the surface of the precambrian basement forms a broad, dissected dome with maximum altitudes around 2200 m east– north-east of ubekendt ejland, from where the surface slopes c. 1°nw and reaches sea level north of innerit and qeqertaq. this dome structure was first observed by rink (1852, p. 20). from its maximum, the dome surface also slopes eastwards down to about 1500 m altitude before it disappears beneath the greenland ice sheet (fig. 5; pulvertaft & larsen 2002). the prevolcanic surface is uneven and constitutes a pre-paleocene etch surface with irregular hilly relief (fig. 5; pulvertaft & larsen 2002; bonow 2005; bonow et al. 2006), which has at least partly been covered by the basalts. the highest preserved basalt cover is at c. 1950 m altitude in the umiammakku sermia – kangilleq area (fig. 5; pulvertaft & larsen 2002). any original basalt cover in the central dome area with altitudes above 1700–1950 m has been removed by erosion. the south-western part of the area considered here, i.e. the northern nuussuaq basin area, comprises much of svartenhuk halvø and is dominated by volcanic rocks, which rest on cretaceous–paleocene sediments of the nuussuaq basin (figs 2, 6). the prevolcanic sediments are exposed in the eastern part of svartenhuk halvø and have been described in overview by larsen & pulvertaft (2000). dam et al. (2009) referred them to the albian–cenomanian upernivik næs formation (deltaic sandstones interbedded with mudstones), the turonian–campanian itilli formation (marine mudstones), and the maastrichtian–danian kangilia formation (mudstones with conglomerates). dam & sønderholm (2021) reviewed the tectonostratigraphic evolution of the nuussuaq basin during the cretaceous and palaeogene with five phases of repeated rifting and faulting with concomitant sedimentation and finally breakup and volcanism. the oldest exposed volcanic rocks in the svartenhuk halvø area are silicic basalts and picrites of the vaigat formation, which are correlated with the upper part of the vaigat formation on nuussuaq and disko (fig. 3). older parts of the vaigat formation may be present below exposure level to the west and to the south of svartenhuk halvø, but it is also likely that the volcanism of the vaigat formation spread gradually northwards from nuussuaq. in any case, volcanism in the nuussuaq basin was well on its way at the time when the oldest volcanic rocks now exposed on svartenhuk halvø were deposited. offshore, the volcanic succession continues on the shelf south, west and north-west of svartenhuk halvø (fig. 1) where seismic investigations have shown the presence of a volcanic package up to several kilometres thick (whittaker 1996; skaarup 2002; skaarup & pulvertaft 2007; gregersen et al. 2013, 2019). gregersen et al. (2013, fig. 5) pointed out a close connection between the onshore and offshore volcanic successions, and drilling and dredging have provided sample material. basalt samples from the 2017 m volcanic succession in the drill hole alpha-1 (location in fig. 1) are compositionally similar to basalts from the svartenhuk formation (l.m. larsen, unpublished data, 2011), and samples from the 1137 m volcanic succession in the drill hole delta-1 (location in fig. 1) are both compositionally and age-wise similar to basalts from the naqerloq formation (nelson et al. 2015). samples dredged on the upernavik escarpment (fig. 1) are compositionally similar to basalts from the svartenhuk formation (polteau & planke 2008). http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 8 of 121 ? ? ? ? ? in ng ia is br æ 2000 m 1000 m 0 m upernavik isfjord laksefjorden qeqertaq fladø innerit upernavik c bf kangilleq karrat isfjord ukkusissat fjord su llu a svartenhuk halvø inngia fj ord rink isbræ upernivik ø ukkusissat illorsuit upernavik kujalleq ubekendt ejland 50 km 1500 m 1000 m 500 m 500 m 500 m 0 m sample 566505 at 1740 m u m ia m m ak ku serm ia 1840 1750 1922 1740 2150 2115 2280 2220 a b 2 1 1 3 gabbro-granite intrusion naqerloq formation svartenhuk formation cretaceous–paleocene sediments vaigat formation precambrian cbf established strike and dip of lavas town, settlement contour lines of prebasaltic surface faults and lineaments with downthrown side indicated cbf proposed (two possibilities a and b) ■■ ■ ■ ■ ■ ■ ■ ■■ ■■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ kangersuatsiaq prøven nuugaatsiaq 71° 72° 72°30' 71°30' 72° 71°53° 52° 51°54° 54° 53° 52°55°56° gv01_02_121_01_lmlfig. 5 geological sketch map of west greenland between 71° and 73°n. the northern and eastern extent of the volcanic rocks evidently continues farther to the north-east beneath the greenland ice sheet. the location of the easternmost basalt sample 566505 is shown. the north-western extension of the cretaceous boundary fault system (cbf) is uncertain and two possibilities are shown, which may both be part of the system. red contour lines highlight the domal structure and uneven character of the prebasaltic surface; the red numbers are altitudes in m a.s.l.. the lines were constructed from the many closely spaced measurements of the altitude of this surface by pulvertaft & larsen (2002, plate 1) and, in the southern part, from the height of the basement peaks. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 9 of 121 structures basin area in the basin south-west of the cretaceous boundary fault system, the sedimentary-volcanic succession is heavily faulted, as described and analysed in detail by larsen & pulvertaft (2000). prevolcanic faulting, uplift, downthrow and erosion in the basin area and along the boundary fault system is indicated by the fact that the volcanic rocks were deposited onto sediments ranging in age from turonian to paleocene; moreover, the sediments contain conglomerates and fanglomerates with basement clasts, some of which are cyclic. locally the sediments show folding, which may be caused by gravity slumping or perhaps postvolcanic faulting (larsen & pulvertaft 2000, p. 12). contemporaneous, interdigitating sediments and volcanic rocks are only found in a few places on south-eastern svartenhuk halvø (larsen & pulvertaft 2000, p. 12–13). synand particularly postvolcanic faulting is evident almost everywhere in the basin on svartenhuk halvø. numerous postvolcanic extensional faults trending nw–se have divided the area into a number of rotated fault blocks with sw-dipping strata with different dips (fig. 6; larsen & pulvertaft 2000, figs 2, 9 and p. 23). dips are smallest in the north-eastern and northern areas (2°–10°) and increase to 10°–45° south of some wnw–ese-trending transfer faults. the largest fault is the nw–se-trending arfertuarsuk fault that runs through the arfertuarsuk inlet in the south to svartenhavn in the north (fig. 6; names on fig. 7) with a downthrow of the north-eastern block of at least 2 km at its south-eastern end (larsen & pulvertaft 2000, p. 24). therefore, the youngest part of the succession is found on the north-eastern side of the arfertuarsuk fault (fig. 6) where the basalts dip 15–23°sw or more, and north thereof where the basalts dip only 0–10°. south-west of the arfertuarsuk fault the vaigat formation reappears at narsinganersua (kap cranstown) together with the older parts of the svartenhuk formation. northern extension of the cretaceous boundary fault system north of svartenhuk halvø, the cretaceous boundary fault system is difficult to follow across the innerit peninsula and qeqertaq because of the largely unfaulted lava pile that has transgressed it. the most easterly possible location of the continuation of the boundary fault north of 72°n is along a se–nw line running from the coast on south-eastern innerit across innerit and qeqertaq, skirting a small exposure of basement at itillia, and continuing out to sea in a nearly straight line (line a in fig. 5). another possibility is that the fault turns south-west in the umiiarfik fjord between svartenhuk halvø and innerit, then turns north-west and crosses southern innerit to southern qeqertaq, then turns north along a fault line there, and then out to sea (line b in fig. 5). this position would explain some of the structures mentioned below (locs d, e and f in fig. 6). the anomaly pattern on the aeromagnetic map of rasmussen (2002) makes a still more westerly position of the cretaceous boundary fault system less likely. northern and eastern areas east of the cretaceous boundary fault system on svartenhuk halvø, the volcanic rocks resting on the high precambrian basement are unfaulted and have a low average dip of 2.3°wnw (larsen & pulvertaft 2000). this is close to the overall slope of the prevolcanic basement surface, which dips 1.5°w to 2.1°wnw between siuteqqut kuuat and kangiusap aaffaa. the maximum preserved thickness is 1300 m beneath the 1700 m high mountain siuteqqut (fig. 7). the volcanic succession north of 72°n is nearly flat-lying and with few exceptions unfaulted. in the western part of the area, the base of the lava pile is below sea level, and to the east, the lavas rest on the precambrian basement that rises from sea level to an altitude of 2000 m close to the greenland ice sheet. the overall dip of the lava succession here is around 1°nw to w, similar to the dip of the pre-basaltic basement surface. in the following we describe some structural features in addition to those south of 72°n, which were treated by larsen & pulvertaft (2000). locs a to g are shown in fig. 6. loc. a. a prebasaltic, n–s-trending fault is presumably situated in sullua fjord east of the innerit peninsula. this is suggested from a jump in surface altitude of the basement across the fjord from 500 m a.s.l. on the eastern side to close to sea level on the western side of the fjord. the lava successions in the profiles 38 and 39 on the opposite fjord sides (fig. 8) roughly correlate across the fjord with the elevation differences caused by the regional dip of 1°w of the lava pile. loc. b. there are indications of synvolcanic subsidence in the southern innerit peninsula just east of the paannivik profile (64, fig. 8), where the lower part of the tunuarsuk member dips 4–6°sw with gradually decreasing dip up-section to 2°sw in the upper part and to 1°sw at the boundary between the nuuit and skalø members on paannivik, see dip measurements on the geological map of larsen & grocott (1991). loc. c. there is a weak anticline on the north-west coast of innerit and the south-east coast of qeqertaq, which could be a primary structure of a low lava shield. loc. d. one of the rare post-basaltic faults in the northern area crosses qeqertaq with a direction close to n–s and a 25 m downthrow to the west. this fault http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 10 of 121 ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ 55° 55° 30 km 54° 71°45´ 72°15´ 71°30´ 72° 56° 54° 72° 71°45´ gv01_02_180_lml.eps 10 200 su llu a su llu a umiiarfik kangiusap imaa umiiviup kangerlua ukkusissat fjord tartuusaq akunnerit saviit ulissat maniiseqqut tasiusap imaa kap cranstown svartenhuk sigguk innerit d d ? ? c c a j e i b f 3 4 2 2 1 3 4 4 4 4 14 9 8 1 5 17 7 9 5 30 13 5 12 10 10 10 15 20 27 40 20 20 20 12 18 7 1 2 3 1 g af af af k l h e qeq er ta q innerit skalø fig. 6 caption and legend on the next page http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 11 of 121 appears to form the boundary between 1°nw-dipping basalts to the east and 1–2°sw-dipping basalts to the west and may reflect the position of the buried cretaceous boundary fault system, which must run out to sea immediately north-west of qeqertaq (fig. 5, line b). the lavas on tukingasoq island also dip 1–2°sw whereas the lavas on kigataq island dip c. 2°w. loc. e. structural analysis of the boundary between the vaigat and svartenhuk formations across umiiarfik fjord indicates uplift (inversion) of the svartenhuk halvø side relative to innerit, suggesting a hidden fault in umiiarfik (larsen & pulvertaft 2000, p. 22). this fault could be part of the cretaceous boundary fault system (fig. 5, line b). loc. f. a concealed nw–se-trending fault crosses southern innerit. this is seen from the change in dip of the lavas across this line from 1°nw to 3°sw southwest of it, indicating a kink in the lava package. nwtrending dykes are common south of the line and rare north of it. this fault could also be part of the cretaceous boundary fault system (fig. 5, line b). loc. g. a postvolcanic fault probably trending nw– se must be present in the sound between innerit and skalø, as indicated by the different dips of the lava packages on innerit (3°sw) and eastern skalø (close to 0°). the lava package on skalø is downthrown c. 500 m relative to that on innerit. this fault may be connected to the large flexure zone that runs through svartenhuk halvø from the south-east coast to the north-west coast (fig. 5), with subsidence to the west of 600 m in northern svartenhuk halvø (larsen & pulvertaft 2000). sv ar te nh uk fo rm at io n skalø member nuuit member va iga t fo rm at io n sv ar te nh uk fo rm at io n tunuarsuk member kuugaartorfik member nunavik member kakilisaat and nerutusoq members prevolcanic sediments (cretaceous to paleocene) precambrian basement covered ice intrusive basalt sheets arfertuarsuk member trachyte flownaqerloq fm monoclinal flexure zone transfer fault/strike-slip fault inferred fault locations described in texta to l arfertuarsuk faultaf major fault with variable sense of movement through time extensional fault, nick on downthrown side ■ ■ ■ ■ ■ ■ gv01_02_182_legende til gv01_02_180_lml.eps fig. 6 (continued) geological map of svartenhuk halvø and the northern volcanic areas. geology south of 72°n after larsen & pulvertaft (2000, fig. 2). north of 72°n, the extent of the volcanic rocks is from escher (1985); the boundary between the tunuarsuk and nuuit members is extrapolated from the existing geological map south of 72°n, combined with the six sampled profiles north of 72°n (see fig. 9), the nearly unfaulted and flat-lying state of the succession with a dip of c. 1°nw and the topographic map. remnants of the skalø member are also present on several summits east of sullua (guarnieri et al. 2022a, b) but are mainly concealed beneath the ice. see fig. 7 for place names mentioned in this study. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 12 of 121 methods field work, sample profiles and map compilation this study is based on geological mapping by j.g. larsen in 1974, 1976, 1980 and 1983, in 1983 together with t.f.d. nielsen, and on geological mapping by k.a. jørgensen in 1981. it is also based on earlier works, particularly by t.c.r. pulvertaft, as described in larsen & pulvertaft (2000). supplementary sampling was done by various field parties of the geological survey of denmark and greenland (geus) in the period 2009–2019. during mapping, the logging and sampling were preferentially done in traverses at locations where the stratigraphy was relatively undisturbed by faulting, although some profiles have been pieced together from shorter sections. flow-by-flow sampling was applied when possible. mapping and location of profiles and samples were done on aerial photographs on a scale of c. 1:40  000, and altitudes were measured with a traditional altimeter. samples from 2009 and later were located and their altitudes measured with gps (global positioning system). in total, 67 numbered and named profiles were measured and described in the field but not all of these were also sampled. the place names mentioned here are shown in fig. 7 and the locations of all profiles are shown in fig. 8. about 1800 samples of mainly lava flows, hyaloclastites, intrusions and sediments were collected. some of the most important and best sampled profiles are designated stratigraphic type and reference sections. for this study, these and some additional profiles mentioned in the text were located on the eox sentinel-2 cloudless satellite image data set available on the greenland portal (https://maps.greenmin.gl/geusmap/?mapname=greenland_portal). for each profile, the utm (universal transverse mercator) coordinates for the start and end points and some intermediate points were read. altitudes were subsequently obtained from the greenland ice mapping project digital elevation model from nasa’s national snow and ice data center (howat et al. 2014). this data set has a resolution of 30 × 30 m, and as most profiles are situated on relatively steep mountain sides, the obtained heights are only moderately accurate. the geographical coordinates and altitudes and the digital elevation map showing the locations of these profiles (fig. 9) were provided by christian brogaard pedersen (geus). their coordinates and approximate altitudes are given in appendix 1 and in digital form in supplementary file s3. the digital elevation model in fig. 9 gives an impression of the topography of the area; of particular notice are the mountainous areas separated by wide valleys with little or no exposure. a low-lying area seen in central western svartenhuk halvø presents a combination of poor exposures, faulting and steep and variable dips of the fault blocks; in this area some geological boundaries are only tentative or not placed at all, e.g. the boundary between the nuuit and skalø members. the most important sampled and analysed profiles are presented as vertical columns in figs 10–13. the heights are m a.s.l., which deviate from true thicknesses in dipping successions; dips and strikes are therefore given below each profile. all the long and well-sampled and analysed profiles are included in the figures, whereas short and unsampled profiles are not. the south and east coasts of svartenhuk halvø present a well exposed section through most of the volcanic stratigraphy in the area. the coastal exposures were photographed from the sea, and the photographs (diapositives) were subsequently projected onto a screen and traced by hand to produce pen drawings of five coastal sections (multi-model photogrammetry was not available at the time). the sections are shown in coloured versions in figs 14–18. their positions are indicated on fig. 8. due to complicated faulting and poor exposures, it was not possible to construct a similar picture of the part of the south coast between akunnerit and saviit. the volcanic parts of the two geological maps at scale 1:100 000 covering svartenhuk halvø were compiled by j.g. larsen using the 1:40 000 vertical aerial photographs, which were also used in the field. the igdlorssuit map sheet was compiled with a sketch master, whereas for the svartenhuk map sheet, the low-lying parts were compiled with a sketch master and the mountanous and difficult areas with a kern pg2 stereo plotter. the precambrian basement areas of the svartenhuk map were compiled in a similar way by j. grocott. geochemistry almost all major-element analyses presented here are by x-ray fluorescence spectrometry (xrf) and were made by jørgen kystol and ib sørensen at the rock geochemical laboratory at the geological survey of greenland (ggu)/geus with analytical procedures as described by kystol & larsen (1999). most elements were determined by xrf on fused glass discs with sodium tetraborate flux; na2o was determined by atomic absorption spectrometry, and feo by potentiometric titration. a subset of samples was analysed for trace elements by jørgen kystol and olga nielsen at geus using a perkinelmer elan 6100 drc quadrupole inductively coupled plasma mass spectrometer (icp-ms). sample disso http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 13 of 121 lution followed a modified version of the procedure used by turner et al. (1999) and ottley et al. (2003). calibration was done using two certified rare-earth element (ree) solutions and three international reference standards with reference values from govindaraju (1994) supplemented with newer values for the ree from the georem database (http://georem.mpch-mainz.gwdg.de). results for reference samples processed and run simultaneously with the unknowns are normally within 5% of the reference value for most elements with concentrations >0.1 ppm. nearly all samples were ground in a tungsten carbide ball mill. this has introduced small amounts of co and ta into the powders. we estimate that 5–15 ppm co has been added to the natural contents of 40–100 ppm co, which does not conceal the natural variations. the amount of ta added is mostly very small but may be significant in the picrites where the natural contents of ta are only about 0.2 ppm. samples with numbers higher than 278607 had major elements analysed at the university of edinburgh. all trace elements were analysed at geus except for samples 572001–572008, which were analysed at actlabs laboratories, vancouver. in total, about 900 samples were analysed for major elements and 300 for trace elements. from the vaigat formation, 166 samples were analysed for major elements and 48 for trace elements. from the svartenhuk and naqerloq formations, 720 samples were analysed for major elements and 233 for trace elements. twenty dyke samples were analysed for major elements and 14 for trace elements. all analytical data are available in supplementary file s4. in the chemical variation diagrams shown here, the elements have been recalculated to 100 wt% on a volatile-free basis. samples with more than 7 wt% volatiles were as a rule not plotted and not analysed for trace elements. all mg-numbers and cipw normative compositions given are calculated with fe2o3/feo adjusted to 0.15. nomenclature the definition of igneous rock types is based on the total-alkalis-silica (tas) chemical classification of le maitre (2002) supplemented by a few extra terms. picrites have mgo ≥12 wt% (le maitre 2002); we use ‘magnesian basalts’ for basalts with 10–12 wt% mgo, which are usually visibly olivine-phyric. ‘silicic basalts’ have slightly increased sio2 = 51–52 wt% and have presumably been slightly crustally contaminated, as uncontaminated basalts in the region rarely have more than 51 wt% sio2. rocks with sio2 >52 wt% are basaltic andesites. we identify rocks, which are contaminated with continental crustal material by their elevated contents of sio2, k2o, rb, ba, th, u, light ree and pb, and low mgo, cao, tio2, feo* (total iron as feo), v, cu and ni relative to common mantle-derived rocks (larsen & pedersen 2009; pedersen et al. 2018). high sio2, k2o and rb and low nb are particularly diagnostic. the term ‘geochemically enriched’ is applied to rocks with elevated contents of a number of incompatible elements such as p, ba, th, u, nb, zr, sr, and the light ree relative to common mantle-derived tholeiitic basaltic rocks. high p, nb and sr are particularly diagnostic. pahoehoe lava flows are commonly thin and highly vesicular, particularly in their upper part. their crust is smooth, with ropy and sometimes glassy top zones. thin pahoehoe flows usually come in groups of flow lobes formed during the same eruption episode, and such groups are called compound lava flows (walker 1971). thick, massive flows with relatively thin and smooth flow tops are here called sheet flows following e.g. self et al. (1997) and sheth (2018). lava flows of aa type are also thick and massive but have thick, highly scoriaceous top zones. on svartenhuk halvø, there is a morphological range between typical pahoehoe sheet flows and aa flows, but typical aa flows are rare. subaqueous lava flows are thick flows that commonly display various kinds of columnar jointing; they are usually brecciated in their upper part and capped by yellowish brown hyaloclastite; their top zones do not show red oxidation. other fully or partly subaqueous flows are pillow lavas with pillows with glassy rims. entablature lava flows have tiers with a basal colonnade with vertical columns overlain by entablature zones showing irregular and fan-shaped columnar jointing. they are formed by emplacement into shallow water. the lithological terms ‘hyaloclastite’ and ‘pillow breccia’ are frequently used synonymously. following white & houghton (2006), we use hyaloclastite as a general term to denote primary clastic subaqueous volcanic deposits. however, we frequently use ‘pillow breccia’ for coarse, clast-rich deposits even though these may include finer-grained clast-poor parts. the term ‘volcaniclastic’ is used as a nongenetic term to denote any clastic deposit with a large component of volcanic material, including both primary volcanic and nonvolcanic deposits such as mass-flows and other redeposited sediments. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 14 of 121 0 10 20 km 72° 72° 54°55° 54°55° 71°45´ 71°30´ 71°45´ 72°15´ 71°30´ qeqertaq marraarnaq tukingasoq kigataq upernavik kujalleq itillia kissaaq kangersuatsiaq prøven qulassivik nuup qaava fladøerne su llu a su llu a qaqqap qaa salliaruseq pannertuup qaqqaa innerit paannivik niaqornaq umiiarfik skalø illukassak nuuit nuuit qaqqaat q ooroq tunuarsuk tunuarsuup kuua peak 1180 n er ut us oq uiarsaariaq amitsup sullua amitsoq serfatm illoorfik innerit am itsup kuua q ooruusaq siggup qoorua maligissap kuua svartenhavn svartenhuk sigguk qaaqqut qaqqaat narsaq akuleqqut kuuat nalluit qaqqaat peak 1050 peak 1091 qiterlikassak oqaasaq taseraarsuit peak 1309 qorlortoq simiutaq uparuaqqusuitsut peak 1211 peak 1270 peak 1260 peak 1140 peak 1625 peak 1370 peak 1829 peak 1430 kissarissoq ukkusissat fjord peak 1250 qinngusaaq qaarsorsuaq simiutaq siuteqqut kuugaartorfik kangiusap imaa itsaku iv is su kk at it in ne ra t peak 1078 firefjeld uiarsaariaq usuit kuussuat qaarajuttorsuaq aputituut akuleqqut nalluarissat qinerfik tasiusaq kussineq arfertuarsuk q innivik tasiusap imaa kap cranstown narsinganersua illerusat qaqqaat tartuusaq akunnerit eq uu tt at k ill iit k uu at peak 1120 saviit kuussuat saviit ulissat qeqertat schade øer qaqqaat qaqqaa ippiup niaqornakassak umiiviup kangerlua peak 1181 issiallak issiallaap qaqqaa ili ve rt oo q kakilisaat kakilisaat maniiseqqut saviit qaqqaat u lissat q ooruat qorlortup kuua siuteqqut kuuat ka ng iu sa p simiuttap kuua in ne rit kuua gv01_04_23_lml aputituut qaqqaat peak 1540 fig. 7 place names used in the text. names in red are also profile names; their numbers are shown in fig. 8. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 15 of 121 fig. 8 locations and numbers of all studied profiles through the volcanic succession. profiles with red numbers are mentioned in the text and their names shown in fig. 7. most of these profiles are shown as vertical schematic ‘logs’ in figs 10–13, and their coordinates are listed in appendix 1. 44e 37 43 66 42 38 39 40 65 64 50 52 51 10 2 1 7 29 55 31 23 27 25 24 5a 5 11 12 4 14 13 26 28 22 21 20 19 17 16 18 15 3636b 35a 35 34 33 33a fig. 14 fig. 16 fig. 17 fi g. 18 fig. 15 32 30 9 6 9a 8b 8c 8a 8 1a 1b 48b 48a 48 4944d 44a 44c 47a 47 31a 54 59 61 67 62 60 63 56 5757a 58a 58b58a 56a 4541a 41b a b 0 10 20 km 72° 72° 54°55° 54°55° 71°45´ 72°15´ 71°30´ 71°45´ 71°30´ gv01_04_24_lml http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 16 of 121 fig. 9 digital elevation model (from the greenland ice mapping project) of svartenhuk halvø with locations of the most important profiles such as type and reference sections. dots are measured points along the profile lines, and colours distinguish individual profiles. see text for detailed explanations. coordinates are in appendix 1 and in digital form in supplementary file s3. 47a peak 1625 m 41b skalø south 12 point 470 m 47 peak 1430 m 4 kakilisaat 40 pannertuup qaqqaa 22 akuleqqut 42 marraarnaq 51 tunuarsuk 64b paannivik 14 saviit north 50 umiiarfik south 55 aputituut 52 nuuit qaqqaat 31 aputituut qaqqaat 13 saviit south 38 qaqqap qaa 37 nuup qaava 58 amitsoq 44c kuugaartorfik 9 peak 1078 m 66 saattut (fladøerne) 48b nerutusoq 30 ivissukkat itinnerat 7 peak 1309 m 65 salliaruseq 67 qooruusaq 2019 47a peak 1625 m 61 qooruusaq 41b skalø south 12 point 470 m 47 peak 1430 m 4 kakilisaat 24 peak 1120 m24 peak 1120 m 40 pannertuup qaqqaa 22 akuleqqut 42 marraarnaq 51 tunuarsuk 10 taseraarsuit10 taseraarsuit 64b paannivik 14 saviit north 50 umiiarfik south 55 aputituut 52 nuuit qaqqaat 31 aputituut qaqqaat 30 ivissukkat itinnerat 64a paannivik64a paannivik 13 saviit south 38 qaqqap qaa 36b arfertuarsuk 60 qooruusaq 29 peak 1050 m29 peak 1050 m 37 nuup qaava 59 qooruusaq 58 amitsoq 62 sigguk 44c kuugaartorfik 9 peak 1078 m 7 peak 1309 m 66 saattut (fladøerne) 48b nerutusoq 0 10 20 30 km 72° 72° 55° 54° 54° 55°56° 71°45´71°45´ 71°30´ gv01_04_22_lml 65 salliaruseq http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 17 of 121 fig. 10 profiles through the volcanic succession in southern svartenhuk halvø. vertical scale is m a.s.l. as measured. dip and strike of the succession are shown below each profile. legend in fig. 13. locations in fig. 9 and coordinates listed in appendix 1. 550 600 650 450 500 250 4 300 350 400 300 350 400 24 peak 1120 m 22 akuleqqut 14 saviit n 950 1000 700 750 800 850 900 200 250 100 50 0 150 300 350 400 550 600 450 500 650 700 750 800 850 900 950 1000 550 600 450 500 650 700 750 800 850 900 400 950 550 600 450 500 650 700 200 250 100 150 300 350 400 550 650 600 450 500 200 250 100 150 300 350 400 181067 164998 164999 165000 165201 165202 165203 165204 165205 165206 sill dyke dyke dyke zone tuff, sed dyke 165211 165210 165209 165208 165207 181066 181065 181064 181063 181062 181061 181060 181059 181058 181057 181056 181055 βo2 181054 181080 181079 181078 sed 181077 181076 181075 e 181074 181073 181023 181022 181021a 181020-21 181019 hy hy hy 181018 181017 181016 e 181014 181015 181013 181012 181011 181010 181009 hy 181008 181081 181082 181083 181084 181085 181086 181087 181025 181026 181027 181028 tuff, sed tuff, sed 181029 181030 mudstone 181031 ? 181032 181033 181034 181035 hy 181036 4 kakilisaat 13 saviit s 12 point 470 m tuff, sed? tuff? tuffs? fault zone 15 16 9 5–1017 gv04_06_040_lml nn72°n 72°n 71°30’n 71°30’n 53°w54°w55°w56°w 25 km 62 13 14 12 4 24 930 55 51 36b 41b 52 37 38 42 40 65 47 10 44c 50 64b 64a 59-61 31 22 http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 18 of 121 fig. 11 profiles through the volcanic succession in central and western svartenhuk halvø. vertical scale is m a.s.l. as measured except for profiles 36b and 59–61. dip and strike of the succession are shown below each profile. legend in fig. 13. index map in fig. 10. locations in fig. 9 and coordinates listed in appendix 1. 550 600 650 450 500 200 250 100 150 300 350 0 50 400 200 250 100 150 300 350 0 50 400 36b arfertuarsuk 59-61 qooruusaq 61b 61a 60 59 59 62 sigguk 650 700 550 600 750 800 850 400 450 500 55 aputituut 1050 950 1000 1050 950 1100 1000 1200 1150 700 profiles 36b and 59–61 are constructed. the vertical scale is approximate thickness, 1 cm = 50 m. 750 600 650 800 850 500 450 550 900 31 aputituut qaqqaat 30 ivissukkat itinnerat 550 600 450 500 650 700 750 800 850 900 950 1000 1050 200 250 100 150 west east foresets hy hy 300 350 0 50 400 550 600 450 500 650 700 750 800 850 900 950 1000 400 9 peak 1078 m 262697 263518 278607 278606 278605tuff tuff tuff tuff tuff t t tuff tuff tuff tuff tuff tuff dyke dyke tuff tuff tuff tuff tuff 278604 278603 278602 278601 278600 278599 278598 278597 278596 dated 278595 278594 278593 278592 278590 278589 278588 278587 278583 trachyte 278581 278586 hy sed 278582 constructed profile 262861 262860 262859 262858 263519 263520 263517 263516 263515 263514 263513 263512 263511 263510 263509 263508 263507 263506 263505 263504 263503 263502 263501 262999 262996 262998 262995 low ti 262994 dyke 262992 262993 262991 262990 262989 262988 262987 262986 262985 262984 262981–82 262983 262977–80 262974–75 262976 262970–71 262973 262972 262969 262968 262967 262966 262964 262965 dyke 263521 263523 24 dyke 263522 dyke 263525 263526 263527 263528 263529 263530 263531 262698 262699 262700 262701 262702 262704 262703 262705 262706 262707 262708 262709 262710 262711 262712 262713 c, e 262714 262714b 262716 262717 c 262718 262719 262720 262721 262722 262723 262724 262725 262726 262738 262739 262741 262742 165238 165239 165242 dyke 165241 165240 profile shift 1200 m actual elevation 165243 165244 165236 glass-rich tuff 165235 165234 165233 welded spatter 165232 165237 251515 251516 251517 251518 tuff 251519 tuff 251520 251521 251522 251523 251524 251525 251526 βo1 β3.1 251527 e 251528 251577 251578 251579 251581 fault 251582 251583 251584 251585 251586 251576 251575 251574 251573 251572 251571 251570 251567 251566 251565 251564 251587 251588e 251589 251590 251568 tuff 251569 251529 251530 dyke 251531 251532 251513, 14 251512 251510 251511 dyke 251509 251508 sill pillow lavas cretaceous sediments 251507 251506 251505 262867 262866 262865 262864 262863 262862 262875 262874 262873 262872 262871 262870 262869 262868 262876 262877 262878 262879 262880 262882 262883 262884 262885 262886 262887 262888 262889 262890 262881 1–2 4–7 3–5 13 6–9 gv04_06_041_lml http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 19 of 121 450 200 250 100 150 300 350 0 50 400 600 500 550 450 200 250 150 300 350 400 41b s skalø profile shift profile shift profile shift 900 850 750 800 700 600 650 550 500 650 700 550 600 750 800 850 900 950 1000 1100 1050 300 350 200 250 400 450 100 gv04_06_042_lml 150 500 50 umiiarfik s 1050 1100 950 1000 1150 1200 1250 700 750 600 650 800 850 500 450 400 550 900 51 tunuarsuk 10 taseraarsuit 1300 1000 1050 1100 900 950 1150 1200 800 850 750 1250 700 750 650 600 550 500 450 400 47 peak 1430 m 750 800 650 700 850 900 950 1000 1050 1100 1150 1200 1250 400 450 300 350 500 550 200 250 600 44c kuugaartorfik 278567 tuff 278568 tuff 278569 278570 278579 278578 dated 278577 278576 278575 278574 278573 tuff 278572 tuff 278571 278566 dated 278565 278564 c 278563 278562 278561 278560 278559 e 278557-58 tuff 278556 278555 278554 278553 278552 278551 278550 278543 278542 278541 278540 278538-39 278536-37 tuff 278535 278534 c 278533 278532 278531 278530 278529 tuff 278528 278527 tuff 278526 278525 278524 278523 278522 278521 278520 278519 278518 278517 278516 278515 tuff 278514 278513 tuff 278512 278511 c 278544 tuff 278545 278546 278547 278548 278549 tuff 263825 263826 263827 c 263828 263829 263830 263831 263832 263833 263834 263836 263837 263838 263839 263840 263841 263842 263851 e 263843 263844 263845 263846 263847 263849 263848 263850 263852 263853 263854 263855 263856 263857 263858 263859 263860 263861 263862 exposures east of profile 263591 263590 263588 263589 tuff 263587 263586 263585 263584 263583 263574 263575 263576 263577 tuff 263582 263580-81 tuff 263578-79 263592 tuff 263593 263594 263595 263596 263597 263598 263599 263600 263801 251639 251640 251641 251642 251643 251644 251645 251646 dyke hy etablatures and pillows 278315 278316 278317 278318 278319 278320 sill 278321 278322 278323 278324 278325 278326 278327 278328 278329 278330 278331 278332 278333 278334 278335 278336 278337 278338 agglomerate 278339 278340 278341 hy hy 278342 278343 278344 278345 278346 278347 coal coal coal 278348 278349 278350 278351 278352 278353 278354 278355 52 nuuit qaqqaat 4 3 0.8 4 10 262899 262900 c 262901 262904 262906 tuff 262905 262907 c 262908 262909 c 262910 262911 262912 262913 262894 β3.1marker horizon dark marker horizon 262902-03 262898 c 262897 262896 262895 262918 262919 262920 278422 278421 278420 278419 278418 278417 278415–16 tuff 278414 278413 278412 278409 278408 278407 278406 278405 pillow lava 278404 278403 278402 c 278401 c 278400 278396–98 278395 278394 278399 dyke 278393 agglomerate 278392 278391 278390 278389 278388 fault zone 278387 278386 278385 278384 278383 278382 278381 278380 278379 278410–11 262914 262915 262916 low ti 251371 251373 251372 dated fig. 12 profiles through the volcanic succession in northern svartenhuk halvø. vertical scale is m a.s.l. as measured. dip and strike of the succession are shown below each profile. legend in fig. 13. index map in fig. 10. locations in fig. 9 and coordinates listed in appendix 1. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 20 of 121 arfertuarsuk mb legend sv ar te nh uk fo rm at io n va ig at fo rm at io n skalø mb nuuit mb naqerloq fm tunuarsuk mb kuugaartorfik mb nunavik mb nerutusoq mb kakilisaat mb mudstone hyaloclastite colonnnade tuff thin-bedded compound flows not exposed gneiss 251278 sample. e: geochemically enriched c: crustally contaminated 200 250 100 150 300 0 50 1050 1100 950 1000 1150 1200 1250 1300 700 750 600 650 800 850 550 64a 64b 900 64a+b paannivik 262793,95 262794 262792 262791 262790 262789 262788 262787 251316-17 251308 sed 251309 251315 251314 251313 entablaturue flow 251312 251311 251310 262786 262785 262784 262783 262782 262781 262780 262779 262778 262777 262776 262775 262774 262773 dated 262772 262771 snow 262770 262769 262768 262767 262766 262765 262764 262763 262762 262761 262760 262759 262758 262756 262755 262754 550 450 500 350 400 300 37 nuup qaava 42 marraarnaq 64b paannivik 650 700 600 750 800 850 900 50 0 500 300 350 200 250 400 450 100 150 300 350 200 250 400 100 50 0 150 251294 skalø mb observed at distance 251293 251292 251291 251290 251289 251288 251287 251286 251285 251284 251283 251282 e 251281 251280 e 251279 251278 251277 251276 e 251275 251273 251274 volc. sed 251271–72 251298 e 251295 ? hyaloclastite mass-flow deposit? pillow lava volcanogenic sediment 38 qaqqap qaa 950 1000 1050 700 750 600 650 800 850 500 550 450 900 700 750 600 650 800 850 300 350 200 250 400 100 50 150 500 550 450 181094 566511 566512 566513 566514 566515 566516 566517 566518 566519 566520 566521 566522 572001 572004 572005 e 572006 e 572007 e 572008 e 181096 181097 181098 181099 181100 181101 foreset -bedded hyaloclastite entablature flow entablature flow 181102 e gneiss gneiss unsampled interval 1050 1100 950 1000 n s 1150 1200 1250 1300 1350 800 850 900 1050 1100 950 1000 1150 900 1400 40 pannertuup qaqqaa 65 salliaruseq 450 500 550 2 4 1 262818 262819 262820 262817 262816 262815e 262814 262813 262812 262811 262810 262809 262808 262807 262806 262805 262804 262803 262802 262801 262800 e 262799 dyke 262798 e 262797 262796 251350 251349 251351 251352 251348 251347 251346 251342 251343 251341 251340 251339 251338 e 251337 251336 black 251335 251334 black 251332 251333 e 251331 black 251330 e 251329 251328 251327 251326 251325 251324 251344 251345 pyroclastic deposit gv04_06_043_lml fig. 13 profiles through the volcanic succession north of svartenhuk halvø. vertical scale is m a.s.l. as measured. in all profiles except paannivik, the succession is near-horizontal, dipping 2° or less. index map in fig. 10. locations in fig. 9 and coordinates listed in appendix 1. the legend applies to figs 10–13. definitions of ‘crustally contaminated’ and ‘geochemically enriched’ in the nomenclature section. volc. sed.: volcanogenic sediment. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 21 of 121 point 546 m km wsw vaigat fmsvartenhuk fm narsinganersua / kap cranstown ene kyst 1 gv03_03_074_01_sh14_lml fig. 14 the coastal section at narsinganersua (kap cranstown). location in fig. 8. the faulted lava flows dip c. 20°sw. the succession comprises 280 m of the uppermost nunavik member of the vaigat formation and a similar thickness of the lowermost tunuarsuk member of the svartenhuk formation. solid black lines denote groups of massive picrite flows that highlight the complicated fault structure. the dotted line is a tuff layer. the apparent low-angle fault in the centre of the section is a wnw–ese-trending fault that cuts obliquely through the exposure (fig. 6); it is partly intruded by a dyke. vertically hatched flows are massive, plagioclase-phyric flows: the lowest flows in the tunuarsuk member and a single basalt flow in the nunavik member. the thickest dykes are indicated by ds. hand traced from colour slides projected onto a screen and therefore not with constant horizontal and vertical scales. see also photo in fig. 34. fig. 15 section along the north coast of tasiusap imaa, 15–20 km north-east of kap cranstown and east of the large fault in arfertuarsuk fjord. location in fig. 8. a: western part. b: eastern part. all three volcanic members of the svartenhuk formation are present. the faulted lava flows dip c. 20°wsw. groups of thin pahoehoe flows are shown with bed-parallel hatching, and tuff layers are dotted. note the 20 m thick tuff at the boundary between the tunuarsuk and nuuit members. the thickest dykes are indicated by ds. samples and observations along the shore are indicated; analyses are in supplementary file s4. hand traced from colour slides projected onto a screen and therefore not with constant horizontal and vertical scales. m skalø member w 6 tuff 6 6 8 6 6 6 tuff tuffs 6 6 6 tuff 6 tuff 6 tuff 6 dyke nuuit member kyst 2 e 8 m tuff 6 6 m glassy welded tuff fault breccia 6 86 6 868 6 86 6 866 tunuarsuk member gv03_03_075_01_sh15_lml a 295 m 45 m310 m b m skalø member w 6 tuff 6 6 8 6 6 6 tuff tuffs 6 6 6 tuff 6 tuff 6 tuff 6 dyke nuuit member kyst 2 e 8 m tuff 6 6 m glassy welded tuff fault breccia 6 86 6 868 6 86 6 866 tunuarsuk member gv03_03_075_01_sh15_lml a 295 m 45 m310 m b http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 22 of 121 fig. 16 section along the south coast of svartenhuk halvø from illerusat qaqqaat to akunnerit. location in fig. 8. the four panels should be read from west to east in the succession a, b, c and d. the faulted lava flows dip 30–40°sw. c and d show c. 2 km of unfaulted succession between the iceberg and akunnerit, which is the thickest unfaulted succession encountered in the vaigat formation, without exposed top and base. groups of thin pahoehoe flows are shown with different bed-parallel hatchings, massive picrite flows are black and massive basalt flows vertically hatched. βo2 is a mapped basalt marker flow near the top of the vaigat formation. note the thick shale and tuff at the boundary between the vaigat and svartenhuk formations. the thickest dykes are indicated by ds. ls: landslide. samples along the shore are indicated; analyses are in supplementary file s4. hand traced from colour slides projected onto a screen and therefore not with constant horizontal and vertical scales. www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 22 of 121 ls ls 6 86 6 8 8 (slipped) 6 8 6 1 m tuff 15 m dark tuff 8 6 8 6 8 6 8 illerusat qaqqaatw km 6 8 6 8 dyke 6 8 6 8 tartuusap qaqqai two massive olivine-rich basalt flows tuff iceberg (naqerloq fm) kyst 3 tartuusaq 6 8 6 8 6 8 6 8 164838possible repetition 15 m shale & tuff βo2 basalt marker flow massive picrite flow 250 m stratigraphy 6 8 tunuarsuk member svarte nhuk formation nunavik member vaigat formation akunnerit thick, massive picrite flows e gv03_03_076_01_sh16_lml point 516 m point 510 m point 544 m a b c d ls ls 6 86 6 8 8 (slipped) 6 8 6 1 m tuff 15 m dark tuff 8 6 8 6 8 6 8 illerusat qaqqaatw km 6 8 6 8 dyke 6 8 6 8 tartuusap qaqqai two massive olivine-rich basalt flows tuff iceberg (naqerloq fm) kyst 3 tartuusaq 6 8 6 8 6 8 6 8 164838possible repetition 15 m shale & tuff βo2 basalt marker flow massive picrite flow 250 m stratigraphy 6 8 tunuarsuk member svarte nhuk formation nunavik member vaigat formation akunnerit thick, massive picrite flows e gv03_03_076_01_sh16_lml point 516 m point 510 m point 544 m a b c d fig. 16 section along the south coast of svartenhuk halvø from illerusat qaqqaat to akunnerit. location in fig. 8. the four panels should be read from west to east in the succession a, b, c and d. the faulted lava flows dip 30–40°sw. c and d show c. 2 km of unfaulted succession between the iceberg and akunnerit, which is the thickest unfaulted succession encountered in the vaigat formation, without exposed top and base. groups of thin pahoehoe flows are shown with different bed-parallel hatchings, massive picrite flows are black and massive basalt flows vertically hatched. βo2 is a mapped basalt marker flow near the top of the vaigat formation. note the thick shale and tuff at the boundary between the vaigat and svartenhuk formations. the thickest dykes are indicated by ds. ls: landslide. samples along the shore are indicated; analyses are in supplementary file s4. hand traced from colour slides projected onto a screen and therefore not with constant horizontal and vertical scales. ls ls 6 86 6 8 8 (slipped) 6 8 6 1 m tuff 15 m dark tuff 8 6 8 6 8 6 8 illerusat qaqqaatw km 6 8 6 8 dyke 6 8 6 8 tartuusap qaqqai two massive olivine-rich basalt flows tuff iceberg (naqerloq fm) kyst 3 tartuusaq 6 8 6 8 6 8 6 8 164838possible repetition 15 m shale & tuff βo2 basalt marker flow massive picrite flow 250 m stratigraphy 6 8 tunuarsuk member svarte nhuk formation nunavik member vaigat formation akunnerit thick, massive picrite flows e gv03_03_076_01_sh16_lml point 516 m point 510 m point 544 m a b c d ls ls 6 86 6 8 8 (slipped) 6 8 6 1 m tuff 15 m dark tuff 8 6 8 6 8 6 8 illerusat qaqqaatw km 6 8 6 8 dyke 6 8 6 8 tartuusap qaqqai two massive olivine-rich basalt flows tuff iceberg (naqerloq fm) kyst 3 tartuusaq 6 8 6 8 6 8 6 8 164838possible repetition 15 m shale & tuff βo2 basalt marker flow massive picrite flow 250 m stratigraphy 6 8 tunuarsuk member svarte nhuk formation nunavik member vaigat formation akunnerit thick, massive picrite flows e gv03_03_076_01_sh16_lml point 516 m point 510 m point 544 m a b c d ls ls 6 86 6 8 8 (slipped) 6 8 6 1 m tuff 15 m dark tuff 8 6 8 6 8 6 8 illerusat qaqqaatw km 6 8 6 8 dyke 6 8 6 8 tartuusap qaqqai two massive olivine-rich basalt flows tuff iceberg (naqerloq fm) kyst 3 tartuusaq 6 8 6 8 6 8 6 8 164838possible repetition 15 m shale & tuff βo2 basalt marker flow massive picrite flow 250 m stratigraphy 6 8 tunuarsuk member svarte nhuk formation nunavik member vaigat formation akunnerit thick, massive picrite flows e gv03_03_076_01_sh16_lml point 516 m point 510 m point 544 m a b c d ls ls 6 86 6 8 8 (slipped) 6 8 6 1 m tuff 15 m dark tuff 8 6 8 6 8 6 8 illerusat qaqqaatw km 6 8 6 8 dyke 6 8 6 8 tartuusap qaqqai two massive olivine-rich basalt flows tuff iceberg (naqerloq fm) kyst 3 tartuusaq 6 8 6 8 6 8 6 8 164838possible repetition 15 m shale & tuff βo2 basalt marker flow massive picrite flow 250 m stratigraphy 6 8 tunuarsuk member svarte nhuk formation nunavik member vaigat formation akunnerit thick, massive picrite flows e gv03_03_076_01_sh16_lml point 516 m point 510 m point 544 m a b c d http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 23 of 121 fig. 16 (continued) section along the south coast of svartenhuk halvø from illerusat qaqqaat to akunnerit. location in fig. 8. the four panels should be read from west to east in the succession a, b, c and d. the faulted lava flows dip 30–40°sw. c and d show c. 2 km of unfaulted succession between the iceberg and akunnerit, which is the thickest unfaulted succession encountered in the vaigat formation, without exposed top and base. groups of thin pahoehoe flows are shown with different bed-parallel hatchings, massive picrite flows are black and massive basalt flows vertically hatched. βo2 is a mapped basalt marker flow near the top of the vaigat formation. note the thick shale and tuff at the boundary between the vaigat and svartenhuk formations. the thickest dykes are indicated by ds. ls: landslide. samples along the shore are indicated; analyses are in supplementary file s4. hand traced from colour slides projected onto a screen and therefore not with constant horizontal and vertical scales. www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 23 of 121 ls ls 6 86 6 8 8 (slipped) 6 8 6 1 m tuff 15 m dark tuff 8 6 8 6 8 6 8 illerusat qaqqaatw km 6 8 6 8 dyke 6 8 6 8 tartuusap qaqqai two massive olivine-rich basalt flows tuff iceberg (naqerloq fm) kyst 3 tartuusaq 6 8 6 8 6 8 6 8 164838possible repetition 15 m shale & tuff βo2 basalt marker flow massive picrite flow 250 m stratigraphy 6 8 tunuarsuk member svarte nhuk formation nunavik member vaigat formation akunnerit thick, massive picrite flows e gv03_03_076_01_sh16_lml point 516 m point 510 m point 544 m a b c d ls ls 6 86 6 8 8 (slipped) 6 8 6 1 m tuff 15 m dark tuff 8 6 8 6 8 6 8 illerusat qaqqaatw km 6 8 6 8 dyke 6 8 6 8 tartuusap qaqqai two massive olivine-rich basalt flows tuff iceberg (naqerloq fm) kyst 3 tartuusaq 6 8 6 8 6 8 6 8 164838possible repetition 15 m shale & tuff βo2 basalt marker flow massive picrite flow 250 m stratigraphy 6 8 tunuarsuk member svarte nhuk formation nunavik member vaigat formation akunnerit thick, massive picrite flows e gv03_03_076_01_sh16_lml point 516 m point 510 m point 544 m a b c d fig. 16 (continued) section along the south coast of svartenhuk halvø from illerusat qaqqaat to akunnerit. location in fig. 8. the four panels should be read from west to east in the succession a, b, c and d. the faulted lava flows dip 30–40°sw. c and d show c. 2 km of unfaulted succession between the iceberg and akunnerit, which is the thickest unfaulted succession encountered in the vaigat formation, without exposed top and base. groups of thin pahoehoe flows are shown with different bed-parallel hatchings, massive picrite flows are black and massive basalt flows vertically hatched. βo2 is a mapped basalt marker flow near the top of the vaigat formation. note the thick shale and tuff at the boundary between the vaigat and svartenhuk formations. the thickest dykes are indicated by ds. ls: landslide. samples along the shore are indicated; analyses are in supplementary file s4. hand traced from colour slides projected onto a screen and therefore not with constant horizontal and vertical scales. ls ls 6 86 6 8 8 (slipped) 6 8 6 1 m tuff 15 m dark tuff 8 6 8 6 8 6 8 illerusat qaqqaatw km 6 8 6 8 dyke 6 8 6 8 tartuusap qaqqai two massive olivine-rich basalt flows tuff iceberg (naqerloq fm) kyst 3 tartuusaq 6 8 6 8 6 8 6 8 164838possible repetition 15 m shale & tuff βo2 basalt marker flow massive picrite flow 250 m stratigraphy 6 8 tunuarsuk member svarte nhuk formation nunavik member vaigat formation akunnerit thick, massive picrite flows e gv03_03_076_01_sh16_lml point 516 m point 510 m point 544 m a b c d ls ls 6 86 6 8 8 (slipped) 6 8 6 1 m tuff 15 m dark tuff 8 6 8 6 8 6 8 illerusat qaqqaatw km 6 8 6 8 dyke 6 8 6 8 tartuusap qaqqai two massive olivine-rich basalt flows tuff iceberg (naqerloq fm) kyst 3 tartuusaq 6 8 6 8 6 8 6 8 164838possible repetition 15 m shale & tuff βo2 basalt marker flow massive picrite flow 250 m stratigraphy 6 8 tunuarsuk member svarte nhuk formation nunavik member vaigat formation akunnerit thick, massive picrite flows e gv03_03_076_01_sh16_lml point 516 m point 510 m point 544 m a b c d ls ls 6 86 6 8 8 (slipped) 6 8 6 1 m tuff 15 m dark tuff 8 6 8 6 8 6 8 illerusat qaqqaatw km 6 8 6 8 dyke 6 8 6 8 tartuusap qaqqai two massive olivine-rich basalt flows tuff iceberg (naqerloq fm) kyst 3 tartuusaq 6 8 6 8 6 8 6 8 164838possible repetition 15 m shale & tuff βo2 basalt marker flow massive picrite flow 250 m stratigraphy 6 8 tunuarsuk member svarte nhuk formation nunavik member vaigat formation akunnerit thick, massive picrite flows e gv03_03_076_01_sh16_lml point 516 m point 510 m point 544 m a b c d ls ls 6 86 6 8 8 (slipped) 6 8 6 1 m tuff 15 m dark tuff 8 6 8 6 8 6 8 illerusat qaqqaatw km 6 8 6 8 dyke 6 8 6 8 tartuusap qaqqai two massive olivine-rich basalt flows tuff iceberg (naqerloq fm) kyst 3 tartuusaq 6 8 6 8 6 8 6 8 164838possible repetition 15 m shale & tuff βo2 basalt marker flow massive picrite flow 250 m stratigraphy 6 8 tunuarsuk member svarte nhuk formation nunavik member vaigat formation akunnerit thick, massive picrite flows e gv03_03_076_01_sh16_lml point 516 m point 510 m point 544 m a b c d http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 24 of 121 saviit qaqqat cm red shale 8 8 6 8 8 conglomerate w. large volcanic clasts 8 m shale, sand and conglomerate w. volcanic clasts m sand, shale and conglomerate 8 m fault8 8 8 wsw major fault nnw–sse minor faults wnw–ese kyst 4 km saviit flexure zone saviit ulissat 8 picrite dykes 8 8 88top of kakilisaat mb sill ene major fault gv03_03_077_01_sh17_lml point 884 m a b saviit qaqqat cm red shale 8 8 6 8 8 conglomerate w. large volcanic clasts 8 m shale, sand and conglomerate w. volcanic clasts m sand, shale and conglomerate 8 m fault8 8 8 wsw major fault nnw–sse minor faults wnw–ese kyst 4 km saviit flexure zone saviit ulissat 8 picrite dykes 8 8 88top of kakilisaat mb sill ene major fault gv03_03_077_01_sh17_lml point 884 m a b fig. 17 section along the south coast of svartenhuk halvø from saviit to ulissat. location in fig. 8. a: western part. b: eastern part. the succession comprises the lower part of the nunavik member. the top of the underlying kakilisaat member is just exposed on the shore in b. panel b contains a cross section of the large flexure zone (here called the saviit flexure zone) that runs nnw–sse across the peninsula (fig. 6) and was described in detail by larsen & pulvertaft (2000, p. 27–28). east of the flexure zone, the lava flows are nearly horizontal; within the zone, a series of wnw–ese strike-slip faults with block rotation have resulted in westerly dips increasing towards the west to more than 40°wsw. west of the flexure zone, the lava flows are faulted but dip fairly regularly only 10–15°wsw. the wnw–ese faults are near-vertical but appear here to be inclined towards ene, which is an effect of the topography and the direction of view. the total offset produced by the faulting is impossible to calculate because of the lithological uniformity of the nunavik member and lack of reliable marker horizons. groups of thin pahoehoe flows are shown with bed-parallel hatching, and massive picrite flows are black. the thickest dykes are indicated by ds. samples and observations along the shore are indicated; analyses are in supplementary file s4. hand traced from colour slides projected onto a screen and therefore not with constant horizontal and vertical scales. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 25 of 121 vaigat formation history. the presence of very olivine-rich volcanic rocks in the lower part of the volcanic succession in the west greenland volcanic province has been known since the pioneering work by steenstrup (1883). the olivine-rich succession was formalised as the vaigat formation by hald & pedersen (1975) with distribution throughout the province from disko to svartenhuk halvø, and with its type area on the north coast of disko between kuugannguaq and asuk. the descriptions in the following are only concerned with the vaigat formation in the svartenhuk halvø area. synonymy. the ‘lower basalt series’ of münther (1973); the ‘lower formation’ of larsen (1981a, b) and holm et al. (1993). subdivisions. the vaigat formation on svartenhuk halvø is divided into three members, from bottom to top: kakilisaat member: brown to grey, silica-enriched, crustally contaminated basalts, basaltic andesites and a few picrites. nerutusoq member: brown basalts and magnesian basalts enriched in incompatible elements. nunavik member: grey picrites and magnesian basalts. the kakilisaat and nerutusoq members are combined into one unit on the geological maps because they cannot be visually distinguished from a distance in the field or on photographs. at close range the two members are distinguishable because the kakilisaat member comprises both plagioclase-phyric and aphyric rocks whereas all the nerutusoq member rocks are aphyric. distribution. the vaigat formation is exposed in the southern, eastern and north-eastern parts of svartenhuk halvø and the south-eastern part of the innerit peninsula (figs 6, 19). south-east of svartenhuk halvø the formation extends to the islands schade øer (qeqertat). in the southern part of the peninsula the volcanic pile is faulted and the lavas dip sw (figs 14–17); in its western part the vaigat formation is covered by the svartenhuk formation and disappears below sea level (figs 14, 16). the westernmost exposure is on the southern tip of qinnivik (kap cranstown) where the upper part of the vaigat formation is exposed (fig. 14). the oldest exposed part of the vaigat formation is situated in the eastern basinal area (fig. 18) immediately west of the cretaceous boundary fault system. farther to the east, the formation has overstepped the cretaceous boundary fault system; here it rests on the precambrian basement and is overlain and overstepped by the svartenhuk formation. the vaigat formation presumably extends at depth beneath south-western svartenhuk halvø and beneath the sea to the south and west. seismic sections in the sea south of svartenhuk halvø and west of ubekendt ejland show a thick package of faulted, w-dipping strata interpreted as lava flows of the svartenhuk formation (skaarup & pulvertaft 2007). the volcanic succession is most probably much thicker than the c. 800 m thickness, which is visible above the first sea-bed multiple, making the presence of the vaigat formation at depth entirely possible. type section. the type section is composed of five individual sections situated close to the coasts of the vaigat strait (pedersen et al. 2017, p. 26). reference sections. on svartenhuk halvø, kakilisaat (profile 4) covers the kakilisaat and nerutusoq members and the lowermost nunavik member. peak 1078 m (profile 9) covers the main part of the nunavik member, and peak 1309 m (profile 7) covers the top part of the nunavik member and the boundary to the svartenhuk formation. thickness. in a swathe of land just west of the cretaceous boundary fault system, the vaigat formation is exposed from bottom to top, dips are less than 5° and faulting is minimal so that thicknesses there can be reliably estimated. in the northern and central parts of this area, thicknesses of 950–1250 m are present, increasing southwards from simiuttap kuua to peak 1309 m northwest of firefjeld. farther south the formation is at least 1250 m thick with no upper boundary preserved. at saviit qaqqaat close to the south-east coast the thickness may have been as much as 1500–1700 m, obtained by extrapolation from thicknesses in the northern part of the basin towards the southern area using a slope of 1° of the lava pile (see the later section on marker horizons in the nunavik member). overall, thicknesses within the basin increase from north to south. the thickness of the subaerial part of the lava pile increases southwards from 200–300 m on innerit, through 450 m in north-central svartenhuk halvø (at peak 1250 m), to 600 m in east-central svartenhuk halvø (at peak 1309 m). on the precambrian basement, the thickness is reduced to c. 500 m in the qinngusaaq area, and the formation thins out at a height of 1100 m east of siuteqqut. the greatest thickness of the vaigat formation is presumed to be present in southern svartenhuk halvø. unfortunately, this is an area where faulting and tilting are severe and prohibit the entire thickness of the formation to be measured in any one profile (figs 16, 17). http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 26 of 121 fig. 18 section along the east coast of southern svartenhuk halvø from maniiseqqut in the south to niaqornakassak in the north. location in fig. 8. the four panels should be read from south to north in the succession a, b, c, and d. the succession comprises the kakilisaat and nerutusoq members and the lower part of the nunavik member, all belonging to the vaigat formation. the kakilisaat member contains three hyaloclastite horizons (hy) of which the upper two are foreset-bedded with n-dipping foresets. in the north, the nunavik member contains one or two hyaloclastite horizons, also foreset-bedded with n-dipping foresets. the solid black flow is a massive picrite flow. the vertically hatched flow is a massive basalt flow (βs2) that forms the top of the kakilisaat member. the position of profile 4 (kakilisaat) is indicated; it ascends just north of a fault scarp and continues along the mountain ridge to point 956 m. see also figs 10, 21. the thickest dykes are indicated by ds. ls: landslide. hand traced from colour slides projected onto a screen and therefore not with constant horizontal and vertical scales. www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 30 of 125 km m ls s nunavik member nerutusoq member kakilisaat member m m m 6 m ls ls ls βs2 marker flow iceberg kyst 5 massive picrite flow hy hy kakilisaat qaqqaat kakilisaat maniiseqqut niaqornakassak kakilisaat 8 8 8 6 m m 6 m m ls ls ls ls iceberg nerutusoq mb not identifiedhyhy ? n gv03_03_078_01_sh18_lml point 510 m point 956 m peak 1140 m point 1068 m a b c d km m ls s nunavik member nerutusoq member kakilisaat member m m m 6 m ls ls ls βs2 marker flow iceberg kyst 5 massive picrite flow hy hy kakilisaat qaqqaat kakilisaat maniiseqqut niaqornakassak kakilisaat 8 8 8 6 m m 6 m m ls ls ls ls iceberg nerutusoq mb not identifiedhyhy ? n gv03_03_078_01_sh18_lml point 510 m point 956 m peak 1140 m point 1068 m a b c dfig. 18 section along the east coast of southern svartenhuk halvø from maniiseqqut in the south to niaqornakassak in the north. location in fig. 8. the four panels should be read from south to north in the succession a, b, c, and d. the succession comprises the kakilisaat and nerutusoq members and the lower part of the nunavik member, all belonging to the vaigat formation. the kakilisaat member contains three hyaloclastite horizons (hy) of which the upper two are foreset-bedded with n-dipping foresets. in the north, the nunavik member contains one or two hyaloclastite horizons, also foreset-bedded with n-dipping foresets. the solid black flow is a massive picrite flow. the vertically hatched flow is a massive basalt flow (βs2) that forms the top of the kakilisaat member. the position of profile 4 (kakilisaat) is indicated; it ascends just north of a fault scarp and continues along the mountain ridge to point 956 m. see also figs 10, 21. the thickest dykes are indicated by ds. ls: landslide. hand traced from colour slides projected onto a screen and therefore not with constant horizontal and vertical scales. km m ls s nunavik member nerutusoq member kakilisaat member m m m 6 m ls ls ls βs2 marker flow iceberg kyst 5 massive picrite flow hy hy kakilisaat qaqqaat kakilisaat maniiseqqut niaqornakassak kakilisaat 8 8 8 6 m m 6 m m ls ls ls ls iceberg nerutusoq mb not identifiedhyhy ? n gv03_03_078_01_sh18_lml point 510 m point 956 m peak 1140 m point 1068 m a b c d km m ls s nunavik member nerutusoq member kakilisaat member m m m 6 m ls ls ls βs2 marker flow iceberg kyst 5 massive picrite flow hy hy kakilisaat qaqqaat kakilisaat maniiseqqut niaqornakassak kakilisaat 8 8 8 6 m m 6 m m ls ls ls ls iceberg nerutusoq mb not identifiedhyhy ? n gv03_03_078_01_sh18_lml point 510 m point 956 m peak 1140 m point 1068 m a b c d km m ls s nunavik member nerutusoq member kakilisaat member m m m 6 m ls ls ls βs2 marker flow iceberg kyst 5 massive picrite flow hy hy kakilisaat qaqqaat kakilisaat maniiseqqut niaqornakassak kakilisaat 8 8 8 6 m m 6 m m ls ls ls ls iceberg nerutusoq mb not identifiedhyhy ? n gv03_03_078_01_sh18_lml point 510 m point 956 m peak 1140 m point 1068 m a b c d km m ls s nunavik member nerutusoq member kakilisaat member m m m 6 m ls ls ls βs2 marker flow iceberg kyst 5 massive picrite flow hy hy kakilisaat qaqqaat kakilisaat maniiseqqut niaqornakassak kakilisaat 8 8 8 6 m m 6 m m ls ls ls ls iceberg nerutusoq mb not identifiedhyhy ? n gv03_03_078_01_sh18_lml point 510 m point 956 m peak 1140 m point 1068 m a b c d km m ls s nunavik member nerutusoq member kakilisaat member m m m 6 m ls ls ls βs2 marker flow iceberg kyst 5 massive picrite flow hy hy kakilisaat qaqqaat kakilisaat maniiseqqut niaqornakassak kakilisaat 8 8 8 6 m m 6 m m ls ls ls ls iceberg nerutusoq mb not identifiedhyhy ? n gv03_03_078_01_sh18_lml point 510 m point 956 m peak 1140 m point 1068 m a b c d km m ls s nunavik member nerutusoq member kakilisaat member m m m 6 m ls ls ls βs2 marker flow iceberg kyst 5 massive picrite flow hy hy kakilisaat qaqqaat kakilisaat maniiseqqut niaqornakassak kakilisaat 8 8 8 6 m m 6 m m ls ls ls ls iceberg nerutusoq mb not identifiedhyhy ? n gv03_03_078_01_sh18_lml point 510 m point 956 m peak 1140 m point 1068 m a b c d km m ls s nunavik member nerutusoq member kakilisaat member m m m 6 m ls ls ls βs2 marker flow iceberg kyst 5 massive picrite flow hy hy kakilisaat qaqqaat kakilisaat maniiseqqut niaqornakassak kakilisaat 8 8 8 6 m m 6 m m ls ls ls ls iceberg nerutusoq mb not identifiedhyhy ? n gv03_03_078_01_sh18_lml point 510 m point 956 m peak 1140 m point 1068 m a b c d km m ls s nunavik member nerutusoq member kakilisaat member m m m 6 m ls ls ls βs2 marker flow iceberg kyst 5 massive picrite flow hy hy kakilisaat qaqqaat kakilisaat maniiseqqut niaqornakassak kakilisaat 8 8 8 6 m m 6 m m ls ls ls ls iceberg nerutusoq mb not identifiedhyhy ? n gv03_03_078_01_sh18_lml point 510 m point 956 m peak 1140 m point 1068 m a b c d km m ls s nunavik member nerutusoq member kakilisaat member m m m 6 m ls ls ls βs2 marker flow iceberg kyst 5 massive picrite flow hy hy kakilisaat qaqqaat kakilisaat maniiseqqut niaqornakassak kakilisaat 8 8 8 6 m m 6 m m ls ls ls ls iceberg nerutusoq mb not identifiedhyhy ? n gv03_03_078_01_sh18_lml point 510 m point 956 m peak 1140 m point 1068 m a b c d km m ls s nunavik member nerutusoq member kakilisaat member m m m 6 m ls ls ls βs2 marker flow iceberg kyst 5 massive picrite flow hy hy kakilisaat qaqqaat kakilisaat maniiseqqut niaqornakassak kakilisaat 8 8 8 6 m m 6 m m ls ls ls ls iceberg nerutusoq mb not identifiedhyhy ? n gv03_03_078_01_sh18_lml point 510 m point 956 m peak 1140 m point 1068 m a b c d km m ls s nunavik member nerutusoq member kakilisaat member m m m 6 m ls ls ls βs2 marker flow iceberg kyst 5 massive picrite flow hy hy kakilisaat qaqqaat kakilisaat maniiseqqut niaqornakassak kakilisaat 8 8 8 6 m m 6 m m ls ls ls ls iceberg nerutusoq mb not identifiedhyhy ? n gv03_03_078_01_sh18_lml point 510 m point 956 m peak 1140 m point 1068 m a b c d http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 27 of 121 fig. 18 (continued) section along the east coast of southern svartenhuk halvø from maniiseqqut in the south to niaqornakassak in the north. location in fig. 8. the four panels should be read from south to north in the succession a, b, c, and d. the succession comprises the kakilisaat and nerutusoq members and the lower part of the nunavik member, all belonging to the vaigat formation. the kakilisaat member contains three hyaloclastite horizons (hy) of which the upper two are foreset-bedded with n-dipping foresets. in the north, the nunavik member contains one or two hyaloclastite horizons, also foreset-bedded with n-dipping foresets. the solid black flow is a massive picrite flow. the vertically hatched flow is a massive basalt flow (βs2) that forms the top of the kakilisaat member. the position of profile 4 (kakilisaat) is indicated; it ascends just north of a fault scarp and continues along the mountain ridge to point 956 m. see also figs 10, 21. the thickest dykes are indicated by ds. ls: landslide. hand traced from colour slides projected onto a screen and therefore not with constant horizontal and vertical scales. www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 23 of 121 fig. 16 (continued) section along the south coast of svartenhuk halvø from illerusat qaqqaat to akunnerit. location in fig. 8. the four panels should be read from west to east in the succession a, b, c and d. the faulted lava flows dip 30–40°sw. c and d show c. 2 km of unfaulted succession between the iceberg and akunnerit, which is the thickest unfaulted succession encountered in the vaigat formation, without exposed top and base. groups of thin pahoehoe flows are shown with different bed-parallel hatchings, massive picrite flows are black and massive basalt flows vertically hatched. βo2 is a mapped basalt marker flow near the top of the vaigat formation. note the thick shale and tuff at the boundary between the vaigat and svartenhuk formations. the thickest dykes are indicated by ds. ls: landslide. samples along the shore are indicated; analyses are in supplementary file s4. hand traced from colour slides projected onto a screen and therefore not with constant horizontal and vertical scales. www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 23 of 121 ls ls 6 86 6 8 8 (slipped) 6 8 6 1 m tuff 15 m dark tuff 8 6 8 6 8 6 8 illerusat qaqqaatw km 6 8 6 8 dyke 6 8 6 8 tartuusap qaqqai two massive olivine-rich basalt flows tuff iceberg (naqerloq fm) kyst 3 tartuusaq 6 8 6 8 6 8 6 8 164838possible repetition 15 m shale & tuff βo2 basalt marker flow massive picrite flow 250 m stratigraphy 6 8 tunuarsuk member svarte nhuk formation nunavik member vaigat formation akunnerit thick, massive picrite flows e gv03_03_076_01_sh16_lml point 516 m point 510 m point 544 m a b c d ls ls 6 86 6 8 8 (slipped) 6 8 6 1 m tuff 15 m dark tuff 8 6 8 6 8 6 8 illerusat qaqqaatw km 6 8 6 8 dyke 6 8 6 8 tartuusap qaqqai two massive olivine-rich basalt flows tuff iceberg (naqerloq fm) kyst 3 tartuusaq 6 8 6 8 6 8 6 8 164838possible repetition 15 m shale & tuff βo2 basalt marker flow massive picrite flow 250 m stratigraphy 6 8 tunuarsuk member svarte nhuk formation nunavik member vaigat formation akunnerit thick, massive picrite flows e gv03_03_076_01_sh16_lml point 516 m point 510 m point 544 m a b c d fig. 16 (continued) section along the south coast of svartenhuk halvø from illerusat qaqqaat to akunnerit. location in fig. 8. the four panels should be read from west to east in the succession a, b, c and d. the faulted lava flows dip 30–40°sw. c and d show c. 2 km of unfaulted succession between the iceberg and akunnerit, which is the thickest unfaulted succession encountered in the vaigat formation, without exposed top and base. groups of thin pahoehoe flows are shown with different bed-parallel hatchings, massive picrite flows are black and massive basalt flows vertically hatched. βo2 is a mapped basalt marker flow near the top of the vaigat formation. note the thick shale and tuff at the boundary between the vaigat and svartenhuk formations. the thickest dykes are indicated by ds. ls: landslide. samples along the shore are indicated; analyses are in supplementary file s4. hand traced from colour slides projected onto a screen and therefore not with constant horizontal and vertical scales. ls ls 6 86 6 8 8 (slipped) 6 8 6 1 m tuff 15 m dark tuff 8 6 8 6 8 6 8 illerusat qaqqaatw km 6 8 6 8 dyke 6 8 6 8 tartuusap qaqqai two massive olivine-rich basalt flows tuff iceberg (naqerloq fm) kyst 3 tartuusaq 6 8 6 8 6 8 6 8 164838possible repetition 15 m shale & tuff βo2 basalt marker flow massive picrite flow 250 m stratigraphy 6 8 tunuarsuk member svarte nhuk formation nunavik member vaigat formation akunnerit thick, massive picrite flows e gv03_03_076_01_sh16_lml point 516 m point 510 m point 544 m a b c d ls ls 6 86 6 8 8 (slipped) 6 8 6 1 m tuff 15 m dark tuff 8 6 8 6 8 6 8 illerusat qaqqaatw km 6 8 6 8 dyke 6 8 6 8 tartuusap qaqqai two massive olivine-rich basalt flows tuff iceberg (naqerloq fm) kyst 3 tartuusaq 6 8 6 8 6 8 6 8 164838possible repetition 15 m shale & tuff βo2 basalt marker flow massive picrite flow 250 m stratigraphy 6 8 tunuarsuk member svarte nhuk formation nunavik member vaigat formation akunnerit thick, massive picrite flows e gv03_03_076_01_sh16_lml point 516 m point 510 m point 544 m a b c d ls ls 6 86 6 8 8 (slipped) 6 8 6 1 m tuff 15 m dark tuff 8 6 8 6 8 6 8 illerusat qaqqaatw km 6 8 6 8 dyke 6 8 6 8 tartuusap qaqqai two massive olivine-rich basalt flows tuff iceberg (naqerloq fm) kyst 3 tartuusaq 6 8 6 8 6 8 6 8 164838possible repetition 15 m shale & tuff βo2 basalt marker flow massive picrite flow 250 m stratigraphy 6 8 tunuarsuk member svarte nhuk formation nunavik member vaigat formation akunnerit thick, massive picrite flows e gv03_03_076_01_sh16_lml point 516 m point 510 m point 544 m a b c d ls ls 6 86 6 8 8 (slipped) 6 8 6 1 m tuff 15 m dark tuff 8 6 8 6 8 6 8 illerusat qaqqaatw km 6 8 6 8 dyke 6 8 6 8 tartuusap qaqqai two massive olivine-rich basalt flows tuff iceberg (naqerloq fm) kyst 3 tartuusaq 6 8 6 8 6 8 6 8 164838possible repetition 15 m shale & tuff βo2 basalt marker flow massive picrite flow 250 m stratigraphy 6 8 tunuarsuk member svarte nhuk formation nunavik member vaigat formation akunnerit thick, massive picrite flows e gv03_03_076_01_sh16_lml point 516 m point 510 m point 544 m a b c d km m ls s nunavik member nerutusoq member kakilisaat member m m m 6 m ls ls ls βs2 marker flow iceberg kyst 5 massive picrite flow hy hy kakilisaat qaqqaat kakilisaat maniiseqqut niaqornakassak kakilisaat 8 8 8 6 m m 6 m m ls ls ls ls iceberg nerutusoq mb not identifiedhyhy ? n gv03_03_078_01_sh18_lml point 510 m point 956 m peak 1140 m point 1068 m a b c d km m ls s nunavik member nerutusoq member kakilisaat member m m m 6 m ls ls ls βs2 marker flow iceberg kyst 5 massive picrite flow hy hy kakilisaat qaqqaat kakilisaat maniiseqqut niaqornakassak kakilisaat 8 8 8 6 m m 6 m m ls ls ls ls iceberg nerutusoq mb not identifiedhyhy ? n gv03_03_078_01_sh18_lml point 510 m point 956 m peak 1140 m point 1068 m a b c d km m ls s nunavik member nerutusoq member kakilisaat member m m m 6 m ls ls ls βs2 marker flow iceberg kyst 5 massive picrite flow hy hy kakilisaat qaqqaat kakilisaat maniiseqqut niaqornakassak kakilisaat 8 8 8 6 m m 6 m m ls ls ls ls iceberg nerutusoq mb not identifiedhyhy ? n gv03_03_078_01_sh18_lml point 510 m point 956 m peak 1140 m point 1068 m a b c d km m ls s nunavik member nerutusoq member kakilisaat member m m m 6 m ls ls ls βs2 marker flow iceberg kyst 5 massive picrite flow hy hy kakilisaat qaqqaat kakilisaat maniiseqqut niaqornakassak kakilisaat 8 8 8 6 m m 6 m m ls ls ls ls iceberg nerutusoq mb not identifiedhyhy ? n gv03_03_078_01_sh18_lml point 510 m point 956 m peak 1140 m point 1068 m a b c d km m ls s nunavik member nerutusoq member kakilisaat member m m m 6 m ls ls ls βs2 marker flow iceberg kyst 5 massive picrite flow hy hy kakilisaat qaqqaat kakilisaat maniiseqqut niaqornakassak kakilisaat 8 8 8 6 m m 6 m m ls ls ls ls iceberg nerutusoq mb not identifiedhyhy ? n gv03_03_078_01_sh18_lml point 510 m point 956 m peak 1140 m point 1068 m a b c d km m ls s nunavik member nerutusoq member kakilisaat member m m m 6 m ls ls ls βs2 marker flow iceberg kyst 5 massive picrite flow hy hy kakilisaat qaqqaat kakilisaat maniiseqqut niaqornakassak kakilisaat 8 8 8 6 m m 6 m m ls ls ls ls iceberg nerutusoq mb not identifiedhyhy ? n gv03_03_078_01_sh18_lml point 510 m point 956 m peak 1140 m point 1068 m a b c d km m ls s nunavik member nerutusoq member kakilisaat member m m m 6 m ls ls ls βs2 marker flow iceberg kyst 5 massive picrite flow hy hy kakilisaat qaqqaat kakilisaat maniiseqqut niaqornakassak kakilisaat 8 8 8 6 m m 6 m m ls ls ls ls iceberg nerutusoq mb not identifiedhyhy ? n gv03_03_078_01_sh18_lml point 510 m point 956 m peak 1140 m point 1068 m a b c d http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 28 of 121 sv ar te nh uk f or m at io n sv ar te nh uk s yd ky st sv ar te nh uk f or m at io n sa vi it m an iis eq qu t ka ki lis aa t m em be r km b hy al oc la st ite fl ex ur e zo ne u lis sa t n un av ik m em be r n un av ik m em be r n un av ik m em be r fi g. 1 9 th e so ut h co as t o f s va rt en hu k h al vø b et w ee n sa vi it an d m an iis eq qu t ( lo ca tio n in f ig . 7 ). th e va ig at f or m at io n (k ak ili sa at a nd n un av ik m em be rs ) i s st ro ng ly a nd re pe at ed ly fa ul te d, a nd th e la va s di p up to 4 0° w . d as he d w hi te li ne s: u pp er li m it of h ya lo cl as tit e ho ri zo ns . d ot te d w hi te li ne : u pp er b ou nd ar y of th e ka ki lis aa t m em be r ( k m b) . f ul l w hi te li ne s: b ou nd ar y be tw ee n th e va ig at a nd s va rt en hu k fo rm at io ns . g eo de tic in st itu te o bl iq ue a er ia l p ho to gr ap h 52 6d n /9 46 5. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 29 of 121 the thickest unfaulted succession encountered in the vaigat formation is c. 2 km without exposed top and base, found on the south coast west of akunnerit (fig. 16). despite good exposures along the south coast of the peninsula the crumbling and monotonous character of the lithologies in many places make repetitions due to faulting very difficult to detect. noe-nygaard’s (1942) estimated thickness of more than 10 km, ignoring any repetitions due to faulting, is definitely too high. the estimate was reduced to at least 8 km by pulvertaft & clarke (1966) and to 2.5–3 km by münther (1973). larsen & pulvertaft (2000) estimated a thickness in the order of 4–4.5 km taking the known repetitions into account. until more studies have been made, we can only conclude that the formation is at least up to more than 2000 m and perhaps up to 3000 m thick. lithology. the lower part of the formation comprises brown to grey hyaloclastites and pillow lavas with subordinate lava flows comprising brown, aphyric and plagioclase-phyric, si-enriched, contaminated basalts, magnesian basalts and subordinate picrites of the kakilisaat member, overlain by brown, aphyric, trace-element-enriched basalts of the nerutusoq member (fig. 18). subaerial lava flows dominate at the south coast west of ulissat. in the lowermost part of the formation minor thin, intervolcanic, tuffaceous and coarse volcaniclastic sediments, sand-, siltand mudstones are present, e.g. on uparuaqqusuitsut (larsen 1981a; dam et al. 2020), south of qorlortoq, on firefjeld and in usuit kuussuat. the upper part of the formation is composed of mostly greenish grey picrites and magnesian basalts of the nunavik member. these form hyaloclastites and overlying subaerial pahoehoe lava flows in the eastern part of the basin. boundaries. the lower boundary of the vaigat formation is the base of the volcanic succession, and the basal lithology is most commonly hyaloclastite. within the basinal area the volcanic succession rests on sediments of cretaceous and danian/selandian age (larsen & pulvertaft 2000; dam et al. 2009, p. 30), but contacts are rarely exposed. a conformable relation between paleocene sediments and the volcanic rocks has been observed in a few places. on the eastern slope of firefjeld, the hyaloclastites rest on a paleocene conglomerate that itself overlies cretaceous mudstones (larsen & pulvertaft 2000; dam et al. 2009, pp. 90 and 115). in south-eastern svartenhuk halvø, a small exposure on the western bank of the river in ulissat qooruat, 5.5 km north of ulissat (shown on the geological map of larsen 1983), shows an ochre-coloured quartzofeldspathic sandstone overlain by black to dark green tuff and dark sandstone with molluscs and ophiomorphae burrows indicating a marine environment, overlain by hyaloclastites (larsen & pulvertaft 2000, p. 13). a conformable transition to hyaloclastites also appears to be present on top of itsaku. interbedded mudstones and hyaloclastites occur along the south-west side of the valley of qorlortup kuua and at issiallak. in other exposures within the basin area prevolcanic tectonic uplift and erosion have removed the upper campanian – maastrichtian and lowermost paleocene sediments. east of the cretaceous boundary fault system, almost all mesozoic–paleocene sediments pre-dating the volcanism have been removed as a result of prevolcanic uplift and erosion. here, the upper part of the vaigat formation rests directly on elevated archaean gneisses and proterozoic metasediments with a substantial surface topography. the lowest volcanic rocks are most commonly hyaloclastites. north of kangiusap imaa, the vaigat formation dominantly consists of hyaloclastites up to 270 m thick, situated at altitudes up to more than 1210 m, the highest level for the hyaloclastites in the vaigat formation, caused by postvolcanic uplift. the upper boundary is marked by the first occurrence of brown, olivine-poor basalts of the svartenhuk formation or, in the northern area, the first occurrence of yellow, quartzofeldspathic sand beds with mudstones and coal seams. the boundary is placed at the base of the sediments because these are interbedded with the basal lava flows of the svartenhuk formation and were presumably deposited during an uplift of the basement rocks east of the cretaceous boundary fault system associated with the earliest eruptions of the svartenhuk formation. in the southern area, no erosion has been recorded to indicate a longer break in the volcanism before the eruption of the svartenhuk formation. however, a sediment/tuff horizon is commonly present at the boundary, for example at tartuusaq (fig. 16) and aputituut qaqqaat (profile 31, fig. 11). the highest altitude of the upper boundary is close to 1300 m and is found at peak 1309 m within the basin in central eastern svartenhuk halvø. the same altitude of the upper boundary is found on the basement near qinngusaaq 12 km north-east of peak 1309 m (larsen & grocott 1991). this shows that the inversion of the basinal area noted by münther (1973) and larsen & pulvertaft (2000, p. 22) is differential with the largest altitude differences in the north, decreasing towards the south-east. geological age. paleocene (danian to selandian). a marine sandstone–tuff–mudstone succession underlying the base of the volcanic succession in ulissat qooruat has provided an early paleocene age: the tuffaceous rocks yielded a dinoflagellate cyst assemblage, which http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 30 of 121 belongs to the deflandrea striata zonule (hansen 1977, 1980, p. 92), now the cerodinium striatum interval biozone referred to the upper part of the lower paleocene (middle np4, powell 1992), corresponding to an age of around 62 ma according to gradstein et al. (2004) and subsequent time scale revisions. chauvet et al. (2019) obtained a 40ar/39ar age of 61.08 ± 0.56 ma for a basalt flow from the lowest part of the vaigat formation at the eastern south coast of svartenhuk halvø. the upper age limit of the vaigat formation is not well constrained. it is estimated to be about 60–61 ma based on the radiometric (40ar/39ar) ages for basalts low in the overlying tunuarsuk member of the svartenhuk formation (between 60.3 ± 1.4 ma and 59.4 ± 0.6 ma, larsen et al. 2016) and an intervening hiatus indicated by a sediment horizon in the northern area. radiometric (40ar/39ar) dating of the vaigat formation in the disko–nuussuaq area (storey et al. 1998) yielded ages of 61–62 ma ± 0.5–1 ma, similar to and slightly older than the radiometric age obtained by chauvet et al. (2019) for the vaigat formation on svartenhuk halvø. the entire vaigat formation was probably deposited within a relatively short time interval (storey et al. 1998; pedersen et al. 2002). age constraints from palaeomagnetic results are described in the section on correlation below. correlation. palaeomagnetic investigations have shown that the lower part of the vaigat formation on nuussuaq is normally magnetised in magnetochron c27n (riisager & abrahamsen 1999), the upper boundary of which is placed at 62.2 ma (vandenberghe et al. 2012; speijer et al. 2020). two palaeomagnetic profiles covering almost the entire lava succession of the vaigat formation on svartenhuk halvø show reversed magnetisation (c26r) throughout (riisager et al. 2003, 2004). also the aeromagnetic pattern of the volcanic areas shows a reversed signature (rasmussen 2002). this indicates that the vaigat formation on svartenhuk halvø correlates with the upper part of the vaigat formation in the disko– nuussuaq area as shown in fig. 3. the palaeomagnetic profiles are located on the southern flank of kakilisaat qaqqaat and the eastern side of peak 1309 m (profile 7, fig. 9). kakilisaat member new member history. on the geological map 71 v.1 syd igdlorssuit (larsen 1983) the kakilisaat member is combined with the overlying nerutusoq member and shown as the units βp, βf1 and βf2. on the geological map 71 v.1 nord svartenhuk (larsen & grocott 1991), the kakilisaat and nerutusoq members are shown as the units βp, βp1, βs1 and βs2. see the two maps provided in supplementary files s1 and s2. the name kakilisaat member was used by larsen & pulvertaft (2000) with the same meaning as here, but the member was not formally defined. name. after kakilisaat qaqqaat, a 956 m high mountain on the coast of south-eastern svartenhuk halvø where the member is well developed. distribution. the kakilisaat member is present in south-eastern svartenhuk halvø and the eastern south coast. it is also present on qeqertat (schade øer) southeast of the peninsula. very similar and perhaps correlative rocks occur as far south as the south coast of ubekendt ejland. the member is covered by younger units to the west as a result of the general westerly dip of the lava pile (fig. 20). the member reappears on the south coast west of the river delta at saviit as a result of a major repetition caused by faulting (fig. 6). the extension of the member north of firefjeld is uncertain because, as already mentioned, it cannot be distinguished from the overlying nerutusoq member from a distance. the absence of the capping βs2 lava flow north of firefjeld and at peak 1309 m (larsen & grocott 1991) suggests that the northward progradation of the kakilisaat member flows stopped here, but this does not exclude that subaquous eruption products could be present farther to the north beneath the nerutusoq member. the kakilisaat member may thus be present in the lowermost hyaloclastite unit along the west side of the valley of siuteqqut kuuat. other possible occurrences include small hyaloclastite remnants on the top of itsaku and on the mountain slope north-west of itsaku; chemical data from these areas are not available. nowhere is the kakilisaat member seen to overstep the cretaceous boundary fault system (see also fig. 26). type section. kakilisaat (profile 4, fig. 10) in south-eastern svartenhuk halvø, in a gully on the east side of the 956 m high mountain kakilisaat qaqqaat, between 420 and 820 m altitude. coordinates in appendix 1. the profile follows the northern side of a fault where the base of the member is covered by scree (figs 18, 21). however, an additional 50 m is exposed near the base of the member, resting on cretaceous sediments at 280 m altitude on the east slope of ulissat qooruat 3 km south-west of the fault. reference sections. saviit south (profile 13, fig. 10) west of saviit; this comprises c. 450 m of subaerial lava flows of the kakilisaat member and includes two picritic lava flows at the base of the profile. peak 1078 m (profile 9, fig. 11) south-west of firefjeld below c. 420 m altihttp://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 31 of 121 kakkak ulissat land 19_1181 kak nun 470 m nun s v nun fig. 20 faulted lava flows of the kakilisaat (kak) and nunavik (nun) members of the vaigat formation in southeasternmost svartenhuk halvø, looking north-west (compare fig. 19). the nerutusoq member is presumed to be present but is unsampled (profile 12, fig. 10). the wide valley running across the picture is ulissat qooruat. beyond this, grey crumbling lava flows of the nunavik member dip westward. in the far background, dark lava flows of the svartenhuk formation (s) overlie grey flows of the vaigat formation (v). photo: kristian svennevig. kak la la la la hy hy hy ner nun nun 956 m fig. 21 the three members of the vaigat formation in the east wall of the mountain kakilisaat qaqqaat, south-east svartenhuk halvø, looking northwest. the kakilisaat member (kak) comprises alternating horizons of hyaloclastites (hy) and subaerial lava flows (la), whereas the nerutusoq (ner) and nunavik (nun) members solely comprise subaerial lava flows. foreset-bedding in the hyaloclastites dips north. a fault is indicated by the dashdot line. the type profile for the kakilisaat member runs uphill just north of the fault and then southwards along the plateau to the summit at 956 m. height of exposed mountain wall c. 560 m (400–956 m). note large masses of slipped material at the foot of the wall. photo: kristian svennevig. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 32 of 121 tude; this includes about 200 m of hyaloclastites at the base, resting on cretaceous heterolithic mudstones. coordinates in appendix 1. thickness. a thickness of more than 450 m of lava flows is present along the south coast west of saviit where the base of the member is not exposed. a thickness of 400 m of hyaloclastite and lava flows is found in the type section at kakilisaat qaqqaat; there are more than 400 m of hyaloclastite and lava flows in the northern part of ulissat qooruat, and 300 m or more of hyaloclastite and lava flows west of firefjeld. thicknesses decrease to 270–200 m of hyaloclastite capped by a few lava flows on the northern slopes of issiallaap qaqqaa, to 170 m of hyaloclastite on southern firefjeld and to only 150 m of hyaloclastites and lava flows along the southern side of eastern usuit kuussuat (fig. 7; see also fig. 26). the rather large variation in thickness was at least partly caused by block faulting of the bedrock prior to the extrusion of the member. lithology. the kakilisaat member is composed of hyaloclastites, pillow lavas and subaerial lava flows of thin pahoehoe and thicker sheet flow types. the member is dominated by basalts and basaltic andesites with brownish weathering colours, but there are also a few magnesian basalts and picrites. the rocks are plagioclase ± olivine-phyric and aphyric; a few carry orthopyroxene phenocrysts. the sheet flows have a distinct brown colour, whereas the thin pahoehoe flows are brownish grey to greenish grey for the more mgo-rich types. the low-viscous magmas of the pahoehoe flows gave rise to well-bedded (foreset-bedded) hyaloclastites, whereas the more viscous magmas of the thicker sheet flows produced weakly bedded hyaloclastites and pillow lavas. it is the distinct, brown sheet flows and their associated brown hyaloclastites that make the member identifiable in the field. the hyaloclastite beds range from 40 m to more than 200 m in thickness, whereas the pahoehoe lava sequences may form up to 60 m thick compound flows; the sheet flows are 10–30 m thick. black mudstones up to several metres in thickness are locally interbedded with the basal subaqueous volcanic deposits. sediment xenoliths are present in a picritic hyaloclastite in a gully on the northern west side of the valley of qorlortup kuua at c. 71°46′20″n, 54°20′30″w. fig. 22 the three members of the vaigat formation on the western slope of the mountain kakilisaat qaqqaat, south-east svartenhuk halvø. the lowest exposure is the uppermost hyaloclastite horizon (hy) in the kakilisaat member (kak); the north-dipping foresets showing the direction of progradation are clearly visible. the hyaloclastites are capped by their associated subaerial lava flows (la). the two uppermost flows in the kakilisaat member with white and brown weathering colours (βs2) are recognisable also in fig. 21. the thin brown flows of the nerutusoq member (ner) are poorly exposed. the normally grey nunavik member (nun) contains a massive brown lava flow, which is an mgo-poor picrite. height of exposure c. 500 m. photo: asger ken pedersen. nun nun ner kak dyke la βs2 hy http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 33 of 121 subaerial and subaqueous facies of the member show regional variations, with subaerial lava flows dominating in the south and south-west. towards the north, the subaerial lava flows transform to hyaloclastite facies with a general north-dipping foreset-bedding. thus the lavas flowed northwards, overriding their prograding hyaloclastites, possibly from a topographical high established close to the eruption sites in the southern area and farther south (see fig. 26). the topmost brown lava flows including the βs2 marker flow (larsen & grocott 1991) reached as far north as south-west of firefjeld before they entered the water, and north of this the entire member is in hyaloclastite facies. the basin in the north presumably formed a fault-bounded trough along the eastern part of the basin. along the south side of usuit kuussuat, a hyaloclastite bed prograded eastwards from dry land in the west (larsen 1981a, fig. 14). subaerial and subaqueous lava facies are interbedded in the area of the type profile at kakilisaat qaqqaat (fig. 21). the type profile is composed of three eruptive units each consisting of thick, foreset-bedded pillow breccias overlain by pillow lava and pahoehoe lava flows. each unit reflects an episode of filling of a basin with a water depth of 35–80 m, indicating basin subsidence between each episode. the north-dipping hyaloclastite foresets also here indicate filling from south towards north (fig. 22). upwards in the profile the hyaloclastites decrease in thickness and the proportion of subaerial lava flows increases. towards the west the amount of hyaloclastite also decreases, and west of ulissat the entire succession is subaerial as in the saviit south profile (fig. 10). near maniiseqqut, the south-easternmost point of svartenhuk halvø, and on schade øer 15–25 km south of maniiseqqut, spectacular pillow lavas, pillow breccias and pillow-rich hyaloclastites are exposed and easily accessible along the shores (figs 23–25). the rocks on schade øer show weak but significant oil staining. an oil seep, several localities with oil staining as well as carbonate veins with oil are also found on the south-eastern shores of svartenhuk halvø, particularly near maniiseqqut (christiansen et al. 1998, 2000). it appears that the lithologies of the kakilisaat member are particularly favourable for trapping of migrating oil. sediments, commonly poorly exposed, are found within the member at several localities. several metres of shaly sediments including gritty layers occur in the lower part of the member at issiallak below the brown breccia and on top of an olivine-phyric hyaloclastite (71°35′26″n, 54°14′32″w, c. 175 m a.s.l.). on the north-eastern part of firefjeld between 390 and 510 m a.s.l., three hyaloclastite layers each about 40 m in thickness are interbedded with and overlain by shaly mudstone and intruded by a sill at 510 m a.s.l. fig. 23 pillow lava with close-lying pillows up to 1 m in diameter developed in the lower part of the flow. the middle part shows flow banding and the upper part is rubbly due to abundant vesicles. height of exposure c. 20 m. kakilisaat member, eastern south coast of svartenhuk halvø near maniiseqqut. photo: asger ken pedersen. fig. 24 elongate pillows reminiscent of seals on the shore. kakilisaat member, schade øer, south of eastern svartenhuk halvø. length of hammer 29 cm. photo: asger ken pedersen. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 34 of 121 a 0.5 m thick shale layer with plant remains and intruded by a basalt sill is present within the brown hyaloclastite in the southern south-west side of the valley of qorlortup kuua at c. 300 m a.s.l., and another 0.5 m thick shale layer with a white surface colour is present on top of the brown hyaloclastite at 337 m a.s.l (71°43′53″n, 54°14′35″w). the latter can be followed towards the north-west due to its white surface colour. it may correspond to a 1.5 m thick shale horizon between the brown and the grey hyaloclastites c. 14 km farther south on the northern flank of issiallaap qaqqaa, as reported by s. munck in rosenkrantz et al. (1942, p. 49, where the mountain was misnamed as umîviup qáqai). in the upstream part of the valley of qorlortup kuua, on the south-western valley side, a 5 m thick silty shale layer is present between an aphyric basaltic hyaloclastite with sediment inclusions and an overlying olivine-phyric hyaloclastite with a loose shale layer on top. the top of the basaltic hyaloclastite appears to be chilled against the sediment, indicating it is intrusive (71°46′23″n, 54°19′34″w, c. 165–184 m a.s.l.). on the north-west side of simiutaq by simiuttap kuua two sedimentary beds at 300–400 m a.s.l. were noted but not visited. they may perhaps correlate with those described here, but could also belong to the overlying nerutusoq member. eruption sites. there is clear evidence from the dip directions of the foreset-bedding in the hyaloclastites and the general stratigraphy that the lava flows of the south-eastern area flowed north and east from a central area around ulissat qooruat and usuit kuussuat, indicating that eruption sites should be present in the southern part of the peninsula and probably also farther south (larsen 1981b; fig. 26). eruptions may have occurred over most of the area occupied by the member because dykes with similar appearance and composition are found at several localities within the member. one eruption site was identified on the lower eastern flank of peak 1309 m where numerous closely spaced small dykes or thin lava sheets cut the hyaloclastites (fig. 26; larsen 1981a, fig. 16). a local eruption site somewhere east of peak 1078 m presumably emitted lavas flowing westwards where they changed into hyaloclastite facies south-east of peak 1078 m (profile 9). a dyke in profile 9 with a chemical composition corresponding to that of the kakilisaat member is probably part of the feeder system for these flows (fig. 26). chemistry and chemostratigraphy. the rocks of the kakilisaat member are dominantly silicic basalts and basaltic andesites; a few picrites extend the total mgo range to 6–13 wt% mgo. relative to the normal rocks of the nunavik member most of the rocks have elevated contents of sio2 and k2o and low mgo, cao, tio2 fig. 25 pillow breccia with abundant, up to metre-sized pillows, pillow fragments and lava fragments in a glass-rich matrix. kakilisaat member, eastern south coast of svartenhuk halvø near maniiseqqut. height of the vertical part of the cliff c. 15 m. photo: asger ken pedersen. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 35 of 121 and feo* (total iron as feo). the elevated contents of sio2, k2o and many trace elements suggest that the rocks have been crustally contaminated (see the earlier nomenclature section and the later geochemistry chapter); the presence of sediment xenoliths in the rocks supports this suggestion. representative analyses are given in the geochemistry chapter. no clear evolution trend with time can be seen in the tio2 contents, whereas the mgo content tends to decrease upwards but with the uppermost basalts being more magnesian again. boundaries. the lower boundary is identical to that of the vaigat formation within the basin, and the sediments underlying the kakilisaat member include marine paleocene mudstones and black, tuffaceous material. the earliest volcanic rocks in the region were therefore deposited in a marine basin. the upper boundary is placed at the top of the uppermost of two massive brown marker lava flows underlying the grey picrites along the south-east coast of svartenhuk halvø; in some areas one of these lava flows may have a white colour due to zeolite-filled cracks (fig. 22). these brown lava flows are designated βf1 on the 1:100 000 igdlorssuit map sheet (larsen 1983) and βs2 on the 1:100 000 svartenhuk map sheet (larsen & grocott 1991). note that the contaminated basalts below βf1 are not distinguished from the picrites on the igdlorssuit map sheet. in the area west of saviit other brown lava flows form the top of the member. in the field, distinction between the kakilisaat and nerutusoq members is difficult based on lithology alone. however, the occurrence of interspersed plagioclase-phyric and aphyric rocks is indicative for the kakilisaat member basalts, whereas the nerutusoq member basalts are all aphyric. definitive distinction between the two members requires chemical analyses; the nerutusoq member rocks have lower sio2 and higher cao than the kakilisaat member rocks and different trace-element patterns (see the geochemistry chapter). geological age. paleocene (danian to selandian). the kakilisaat member is reversely magnetised and was erupted during magnetic polarity chron c26r (riisager et al. 2003). its age is thus constrained to the interval 62.2–61 ma, based on the lower age boundary of c26r and the upper age boundary of the vaigat formation. chauvet et al. (2019) presented an 40ar/39ar age of 61.08 ± 0.56 ma in accordance with this for a silicic basalt lava flow from the lower part of the vaigat formation at the eastern south coast of svartenhuk halvø. the location and chemical composition of the basalt (agranier et al. 2019, their site 42) indicate that it belongs to the kakilisaat member. correlation. the kakilisaat member is suggested to be time-correlated with the crustally contaminated tunoqqu and kûgánguaq members of the vaigat formation on nuussuaq and disko (fig. 3). all were erupted within a period when the conditions for high-level magma storage and reaction with crustal rocks were favourable. nerutusoq member new member history. on the geological maps of larsen (1983) and larsen & grocott (1991) the member is combined with the underlying kakilisaat member because the two members can only be visually distinguished at close range. the name nerutusoq member was first used by larsen & pulvertaft (2000) in the same sense as here, but the member was not formally defined. name. after the large, nne–ssw-oriented inland valley nerutusoq in northern svartenhuk halvø (figs 7–9). distribution. the member is present in the northern part of svartenhuk halvø along the valley of nerutusoq (fig. 27; see also larsen & pulvertaft 2000, plate 1, profile c–d) and around the simiutaq hill. it is also present in the short valley that extends w–e from just north of the lake taseraarsuit around 71°47′n, 54°25′w to the valley of siuteqqut kuuat (see also fig. 40). the member is inferred to be present in the valley of siuteqqut kuuat, exposed on the western valley side close to the cretaceous boundary fault system, and also in the broad area around firefjeld. the member is present in south-eastern svartenhuk halvø, where a few flows have been identified in profiles 4 and 13 at kakilisaat and saviit south (fig. 10). the extension of the member to the west, north-west and north is unknown. type section. nerutusoq (profile 48b) in the northern part of svartenhuk halvø along the nerutusoq valley (figs 7–9, 27, coordinates in appendix 1). no vertical profile such as those in figs 10–13 could be constructed because a major part of the profile is near-horizontal, although it is up-section through inclined hyaloclastites. the lower boundary of the member is placed on top of cretaceous sediments c. 100 m a.s.l. at the south-eastern end of simiuttap kuua, just west of the hill simiutaq. the profile follows the eastern riverside in the nerutusoq valley 10.5 km southward until a gully in the western valley side, which is then followed uphill (fig. 28). most of the profile transects sw-dipping hyaloclastites, but in the gully these are capped by brown aphyric basalt flows at 595 m a.s.l. (fig. 28). at 670 m a.s.l. the brown basalts are overlain by thin grey pahoehttp://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 36 of 121 hoe flows of the nunavik member. here the exposure is downwarped along a nw–se fault and dips to the south. on the north-eastern side of the fault the boundary is horizontal with an elevation of 750 m a.s.l.. the lower boundary in the type profile was not visited in the field, and therefore we cannot know if the lower part of the section includes brown hyaloclastites of the kakilisaat member. the lowest analysed sample (ggu 263818) in the type profile is from 360 m altitude in a small drainage stream c. 2.8 km north-west of peak 1250 m. the undifferentiated βp hyaloclastite of the map sheet (larsen & grocott 1991) reflects the uncertainty concerning the field distinction between the kakilisaat and nerutusoq members (see the description of the boundaries of the kakilisaat member). reference sections. taseraarsuit (profile 10) on the south-west side of peak 1140 m, 2–3 km north of lake taseraarsuit (figs 7–9, 12; coordinates in appendix 1). at 350–670 m a.s.l., this profile is mostly well exposed and presents 320 m of the upper part of the member, including 240 m of hyaloclastites, 40 m entablature flows, 40 m fully subaerial, thin lava flows and the boundary to picrite lava flows of the overlying nunavik member. a second reference section is kakilisaat (profile 4) in south-eastern svartenhuk halvø. at 825–860 m a.s.l. the nerutusoq member here has a strongly reduced thickness of only 35 m of subaerial lava flows without a significant break in the volcanism (figs 10, 18). thickness. an estimated maximum thickness of up to 650 m may be present in the southern part of the 55° 55° 10 km 54° 71°45´ 71°30´ 72° 56° 54° 55°56° 54° 55°56° 56° 54° 72° 71°45´ gv01_02_180_01_lml.eps umiiar�k kangiusap imaa arfertuarsuk tartuusaq akunnerit saviit kuussuat ulissat qooruat maniiseqqut schade øer nunavik mb lava �ow directions: itsaku usuit kuussuat aputituut qaqqaat 1309 m fire�eld simiutaq ner ut us oq tasiusap imaa 3 4 2 2 1 3 4 4 4 4 14 9 8 1 5 17 7 9 5 13 5 12 10 10 15 20 27 40 20 20 20 12 18 1 2 3 1 skalø nerutusoq mb kakilisaat mb nunavik mb and tu� layer known eruption sites: nerutusoq mb main eruptive areas: kakilisaat mb kakilisaat mb tunuarsuk mb nuuit mb nerutusoq mb total known extent: kakilisaat mb fig. 26 map showing flow directions as indicated by dips of foreset-bedding in hyaloclastites of the vaigat formation and the main eruptive areas and total known extents of the kakilisaat and nerutusoq members. the coloured area at nerutusoq is the ‘table mountain’ with subaerial lava flows described in the text. both members probably continue below exposure level to the west. the kakilisaat member may also be present farther north than shown. the nunavik member is considered to extend across the entire area. also shown are known eruption sites for the volcanic succession. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 37 of 121 nerutusoq valley, dependent on the absence or presence of the kakilisaat member beneath it and the topography of the bedrock. with the same reservations, the thickness in the northern part of nerutusoq is about 300 m. towards north-west along the south-west side of the valley of simiuttap kuua the thickness decreases from 300 m to 100 m. near lake taseraarsuit the thickness is more than 230 m, and along the west side of the valley of siuteqqut kuuat the member is inferred to be present with a thickness of more than 200 m. the member is strongly reduced in thickness to the south-east and south, where it is 60 m thick at kakilisaat (profile 4) and 40 m thick at saviit south (profile 13). lithology. the nerutusoq member is composed of brown aphyric basalts and magnesian basalts forming hyaloclastites and pillow lava tongues, which in places are capped by an up to 50 m thick succession of subaerial, thin, vesicular pahoehoe lava flows. the lowermost lava flows were emplaced into shallow water and are entablature lavas with irregular and fan-shaped columnar jointing and an uneven top topography; the flow tops have a dark brown colour and are presumably glass-rich. in the southern part of the nerutusoq valley, the member consists of southerly dipping hyaloclastites capped by lava flows at altitudes of 720–750 m. this succession can be followed northwards on the western valley side; 2 km farther north the brown lava flows have all transformed into hyaloclastites, which dip northwards below hyaloclastites of the overlying nunavik member at altitudes of 720–400 m, decreasing northwards. the fig. 27 the vaigat formation west of the siuteqqut kuuat valley, looking west. in the foreground subaerial thin pahoehoe lavas of the nunavik member overlie 300 m of associated hyaloclastites with large-scale foreset-bedding indicating volcanic progradation to the north (enhanced by drawn, black lines). the marker unit βo1 with dark flows is indicated by a white [; the unit disappears into hyaloclastite facies at the position of the white star. in the middle ground is the wide nerutusoq valley. in its western mountainside lavas and hyaloclastites of the nerutusoq member are exposed in a small area on the low slope, overlain by the nunavik member. the dashed black line is the upper boundary of the hyaloclastite facies; the facies boundary crosses the member boundary (dotted black line). in the distant mountains thick-bedded lava successions of the svartenhuk formation overlie the nunavik member above the white line. profiles 50 (umiiarfik south) and 51 (tunuarsuk) are indicated with yellow lines. geodetic institute oblique aerial photograph 524cv/8298. svartenhuk formation nunavik member nunavik member nerutusoq member hyaloclastite hyaloclastite profile 51 profile 50 βo1 http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 38 of 121 hyaloclastites of the two members appear to interdigitate (see cross-section in larsen & pulvertaft 2000, plate 1, profile c–d). the two members were erupted at a time with similar water levels, at least partly contemporaneously, and thus formed hyaloclastites up to the same altitude of 720 m, covered by lava flows. a similar situation exists in the valley of lake taseraarsuit where more than 250 m of east-dipping, brown, aphyric hyaloclastites are interdigitated with grey hyaloclastites of the nunavik member down to 400 m a.s.l. on both sides of the valley (fig. 29). here, the top of the hyaloclastites of the two members reaches an altitude of 600 m. profile 10 in the northern side of the valley has large pillows with south-dipping longitudinal axes. thus, the hyaloclastites and pillow lavas dip away from a common centre located in the southern nerutusoq valley (fig. 26). in the southernmost part of the nerutusoq valley, the member appears to end in a steep cliff against the picrites to the north-west; this may be a secondary feature caused by faulting or landslipped nunavik member lava flows. in the southern part of svartenhuk halvø a few subaerial lava flows of the nerutusoq member were deposited on elevated dry land formed by the kakilisaat member basalts. the reduced thickness of the member in this area may also be caused by its distal position relative to the main source area in the north. eruption sites. as described above, the hyaloclastites and pillow lavas dip away from a common centre located in the southern nerutusoq valley. this was probably a major eruption centre for the member. overall, the major volume of the nerutusoq member forms a nw–seelongated ‘table mountain’ of hyaloclastite capped by cogenetic lava flows, centred on a major eruption area in the nerutusoq valley area (fig. 26). this mountain acted as a barrier for the extension of the lower part of the nunavik member. the barrier presumably terminated in the area north-west of firefjeld and was not important in the south. the situation to the west and north-west is unknown. in the northern nerutusoq valley and along the west side of the valley of siuteqqut kuuat, the brown hyaloclastites show mainly westerly-dipping foreset-bedding and have no capping subaerial lava flows of the member. they were probably fed from a number of nnw– sseto ne–sw-trending, aphyric basalt dykes, which cut the hyaloclastites. several thin, n–s-trending possible feeder dykes are also present in the taseraarsuit valley. an early, widespread eruptive phase through the dykes therefore appears to have taken place before the volcanism was concentrated in the table-mountain area. chemistry and chemostratigraphy. the nerutusoq member consists of basalts and magnesian basalts with fig. 28 the middle part of the type profile (48b, fig. 9) for the nerutusoq member. the location is where the profile leaves the riverbed in the main valley and climbs uphill through a gully in the hyaloclastites (hy) and into the subaerial lava succession (la). western slope of the southern nerutusoq valley, central svartenhuk halvø. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 39 of 121 8–12 wt% mgo, normal sio2 contents of 48–49 wt%, relatively high contents of cao (12–13.6 wt%), and high p2o5 (0.20–0.32 wt%). the p2o5/tio2 ratios are high, 0.116– 0.17. besides p2o5, the member is relatively enriched in a number of trace elements such as ba, th, u, nb, zr, sr and the light rare-earth elements. representative analyses are given in the geochemistry chapter. boundaries. the nerutusoq member rests on lava flows and hyaloclastites of the kakilisaat member and in the valley of simiuttap kuua possibly on sediments of paleocene or cretaceous age. as described above, distinction in the field between the kakilisaat and nerutusoq members is difficult except at close range; they are therefore shown as one unit on the published geological maps. well-defined boundaries to lava flows of the kakilisaat member are exposed at kakilisaat (profile 4) and saviit south (profile 13) in south-eastern svartenhuk halvø (fig. 10); there is no indication of a break in the volcanic activity at the boundary. in the type profile (48b) the lower boundary, which was not visited in the field, is belived to be present along the south-west side of the simiuttap kuua valley, with contact towards either cretaceous sediments or the top of the kakilisaat member, if this is present here. two sedimentary beds between 300 m and 400 m a.s.l. on the north-west side of simiutaq may be close to the lower boundary. in other places the boundary is not exposed or has not been identified. a 0.5 m thick bed of black mudstone within the brown hyaloclastite southeast of peak 1309 m may form the lower boundary towards the kakilisaat member here. the upper boundary is best exposed in the valley sides of nerutusoq in the type profile (48b) and in profile 10. the boundary is distinct because of the colour contrast between the brown lavas of the nerutusoq member and the overlying grey picrites of the nunavik member. in both profiles there is no indication of any break in the volcanism at the upper boundary; rather, the nerutusoq member appears to be partly contemporaneous with the nunavik member. a break in the volcanism may be indicated by a shale layer with plant remains at the boundary between a brown and a grey hyaloclastite 3.8 km east-south-east of peak 1309 m (71°43′53″n, 54°14′35″w, c. 337 m a.s.l.). however, the brown hyaloclastite could also belong to the kakilisaat member. the upper boundary reaches a maximum altitude of 860 m at kakilisaat, 750 m in southern nerutusoq and 600–650 m in the valley of taseraarsuit. geological age. paleocene (danian to selandian). the nerutusoq member is reversely magnetised and was erupted during magnetic polarity chron c26r (riisager et al. 2003). its age is thus constrained to the interval 62.2–61 ma, based on the lower age boundary of c26r and the upper age boundary of the vaigat formation. no reliable radiometric age has been obtained for the member. correlation. none known. nunavik member new member history. on the geological map 71 v.1 syd igdlorssuit (larsen 1983) the nunavik member is shown as part of the unit βp (hyaloclastites) and as the unit βo (lava flows). on the geological map 71 v.1 nord svartenhuk (larsen & grocott 1991) the member is shown as the units βp2 (hyaloclastites) and βo (lava flows); hyaloclastites of the member may also be included in the unit βp (undifferentiated hyaloclastites). fig. 29 brown hyaloclastites (dark) of the nerutusoq member interdigitated with grey hyaloclastites (light) of the nunavik member. both units with east-dipping layering. the boundary between brown and grey units is traced with a thin black pen for clarity. north-eastern corner of the valley with the lake taseraarsuit, bordering to the siuteqqut kuuat valley. height of exposure about 200 m. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 40 of 121 name. after nunavik, the former greenlandic name for svartenhuk halvø (now sigguup nunaa). distribution. the nunavik member is exposed over large areas in the eastern half of svartenhuk halvø. the member dominates the south coast as far west as tartuusaq. the westernmost exposure is a small outlier on south-western svartenhuk halvø at narsinganersua/ kap cranstown. to the north, the member reaches the southern part of the innerit peninsula where the dips of the hyaloclastite bedding indicate mostly northerly flow directions. its extension below exposure level north of 72°00′n may be up to 10 km, based on a northward reduction in thickness (see the section on thickness). a significant extension of the member occurs east of the cretaceous boundary fault system from kuugaartorfik in the north-west to qinngusaaq and uparuaqqusuitsut in the south-east. due north of itsaku, the easternmost extension of the member consists of erosional remnants in two small outliers at peaks 1211 and 1270 m (figs 6, 7), up to 17 km east of the boundary fault. these rocks are situated at some of the highest altitudes for the member due to the postvolcanic uplift discussed by pulvertaft & larsen (2002). altitudes reach c. 1300 m in a small erosional remnant 2.5 km north-east of qinngusaaq where the upper boundary of the member is preserved (larsen & grocott 1991). due to the general westerly dip of the volcanic succession, the lower part of the member is well exposed in the eastern part of the basin including the south coast of svartenhuk halvø (fig. 30). the upper part of the member is exposed towards west and south-west, on the precambrian basement to the east, and on the innerit peninsula (fig. 31) to the north. the entire member is exposed in the nearly flat-lying areas of the central and north-eastern parts of the basin. type section. due to faulting, especially in the southern part of the peninsula, it has not been possible to establish a type section that includes the maximum thickness and entire stratigraphy of the nunavik member. peak 1078 m (profile 9) is designated as type section because it is unfaulted, well-sampled and analysed. the profile runs along the south-east side of peak 1078 m, from 470 m to 1078 m a.s.l. (figs 9, 11; coordinates in appendix 1). the lower 600 m of the member are exposed here, but the top is missing; the base is not exposed but is seen 1–2 km south-west of the profile. reference sections. both base and top of the nunavik member are exposed in the central and northern areas where the thickness is reduced. a profile on the south side of peak 1309 m (profile 7; coordinates in appendix 1) is a reference section; it includes the thickest undisturbed succession in the area, and dips are only 1–2°nnw. the profile was not sampled for geochemical analysis and is not shown in figs 10–11. the upper boundary is well exposed. the position of the lower boundary cannot be determined on the basis of the present knowledge because grey and brown hyaloclastites of unknown composition are intercalated here. the entire profile is c. 1100 m in thickness. it begins at c. 170 m altitude in a river north-west of firefjeld, where cretaceous sediments are overlain by olivine-phyric pillow lavas succeeded up-river by intercalated grey and brown hyaloclastites. at 230 m the profile turns north and follows a side stream uphill; from 450 m grey picritic hyaloclastites belong to the nunavik member. at c. 570 m the grey hyaloclastites are overlain by subaerial lava flows of the nunavik member. from here the profile climbs the hillside to the top at 1309 m (figs 9, 32; see also fig. 40). at the top about 10 m of the lowest flow of plagioclase-phyric basalt of the overlying tunuarsuk member is preserved. a palaeomagnetic sample profile by riisager et al. (2003, 2004) was collected from the subaerial part of this profile. the second reference section is peak 1120 m (profile 24) on the north-western valley side of the headwaters of saviit kuussuat, where the upper 700 m of the member are exposed from 270 m to 970 m altitude (figs 9, 10; coordinates in appendix 1). the profiles cannot be correlated by field observations or chemical data, a fact which applies to most profiles in this member in areas where faulting is common. profile 24 is located west of a flexure zone south of usuit kuussuat, which has caused a downthrow of at least 90 m of the lavas to the west (fig. 33; larsen & pulvertaft 2000, p. 27 and fig. 11). thickness. from south to north in the basinal area, thicknesses decrease from in excess of 930 m (including 120 m hyaloclastite) in the issiallaap qaqqaa area, to 855 m (300 m hyaloclastite) in the profile of peak 1309 m, to 800 m (250 m hyaloclastite) at peak 1250 m south of simiutaq, and 600–900 m (c. 300 m hyaloclastite) on the south-west side of the simiuttap kuua valley. on innerit the exposed thickness is 380 m, including 180 m of hyaloclastite, and a total thickness of around 500 m is possible if the hyaloclastites continue below sea level to attain a thickness of 300 m, similar to that on svartenhuk halvø. the reduced thickness of the subaerial lava succession on innerit corresponds to a lava shield with a northerly slope of 1° (as argued in the description of marker horizon βo1). by extrapolating the decreasing thickness of the subaerial lava pile to the north, it is inferred that the lava flows of the member may have travelled up to 10 km farther north before the uppermost flow had reached sea level and the progradation stopped. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 41 of 121 in the central area around nerutusoq, the thickness is only 325 m because an eruption centre of the underlying nerutusoq member formed a local high (fig. 26) around which the lowest nunavik member lavas flowed. on southern svartenhuk halvø, the western outlier at kap cranstown comprises 280 m of the upper part of the member (fig. 34). along the south coast from tartuusaq to maniiseqqut, uplift along numerous nw– seand wnw–ese-trending faults has exposed the nunavik member for 32 km (fig. 30). it is difficult to estimate the thickness of the member in the southern area due to the faulting and lack of marker horizons. 884 m d d dd d sh028_jgl_c3_31_76mod+jgl_c3_32_76mod d fig. 30 picrites typical of the nunavik member forming a thick succession of grey, crumbling, compound pahoehoe lava flows dipping 15–20°wsw. several dark dykes (d) cut the succession. south coast of svartenhuk halvø between saviit and ulissat; the mountain is saviit qaqqaat (884 m). see also fig. 17. nunavik mb nuuit mb skalø mb la lak k kk hy hy tunuarsuk mb fig. 31 the northernmost exposures of the vaigat formation on the innerit peninsula. the formation here comprises 300–350 m of hyaloclastites (hy) and subaerial picrite lava flows (la) of the nunavik member. these are overlain by the svartenhuk formation, which comprises poorly exposed, whitish quartzofeldspathic sediments of the kuugaartorfik member (k) overlain by lava flows of the tunuarsuk, nuuit and skalø members. south-eastern innerit peninsula beneath the paannivik mountain; compare fig. 46. the highest peak on the skyline is at 1288 m. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 42 of 121 βo1 βo1 βo2 1309 m fig. 33 well-exposed lavas of the nunavik member in the large flexure zone that runs sse from usuit kuussuat to the south coast west of ulissat (fig. 6). over the about 700 m from right to left in the photo, the dip of the lava flows changes from 5°wsw to 30°wsw through two sets of closely spaced faults oriented wnw–sse and nnw–sse (larsen & pulvertaft 2000). displacements on the faults are difficult to see due to the monotonous lithology of the nunavik member; an interpretation is shown in fig. 17 (saviit flexure zone). south coast of svartenhuk halvø just west of ulissat. fig. 32 the nunavik member close to reference profile 7. the succession of thin, grey, subaerial picrite lava flows is c. 690 m thick; the underlying hyaloclastites are exposed in the river gully in the lower part of the photo. the mapped marker horizon of dark flows, βo1, is indicated. the βo2 flow near the top is a widespread basalt marker flow in the nunavik member. the topmost small remnant of a lava flow is the lowest plagioclase-phyric basalt of the tunuarsuk member. peak 1309 m, central-eastern svartenhuk halvø, viewed from the south. the position of profile 7 cannot be shown here but is indicated in fig. 40. the mountain slope seen here is the site of an important palaeomagnetic profile located on a similar photo in riisager et al. (2003, fig. 4). photo: asger ken pedersen. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 43 of 121 the maximum thickness in a single fault block is around 2000 m without the presence of either base or top (fig. 16 east of tartuusaq). thus the thickness of the nunavik member in the southern area is certainly more than 2000 m. calculation of the actual thickness depends on estimates of the throw on many faults and subsidence in flexure zones. possible shingling of the lava pile, i.e. lateral shift of volcanic deposits with time, would be another complicating factor because it makes the stratigraphic thickness different from actual thicknesses at any location. further discussion is found in larsen & pulvertaft (2000, p. 34–35). on the precambrian basement, thicknesses decrease eastwards from about 400–500 m in the west on qinngusaaq and kuugaartorfik, and the member wedges out at an altitude of around 1100 m in the north-east. in the easternmost outliers at peaks 1211 m and 1270 m, 250–270 m of picrites comprising hyaloclastites, pillow lavas and subaerial flows were erupted through local feeder dykes (fig. 26). lithology. the nunavik member is characterised by olivine-rich tholeiitic picrites and olivine-phyric to olivine-microphyric magnesian basalts and subordinate basalts. the member consists of hyaloclastites and coarser pillow breccias overlain by subaerial lava flows of compound pahoehoe type together with a few prominent, thick sheet flows. tuffs and soils are relatively rare, forming red beds 5–20 cm in thickness between the lava flows. other sediments are scarce and include volcanogenic conglomerates (fig. 35) and sandstones, and more fine-grained silty beds. the member has light greenish grey or purplish grey weathering colours. the light grey colour is due to zeolite-filled vesicles. a few erosion-resistant picritic lava flows may have brown weathering colours and may therefore resemble brown basaltic, mgo-poor lava flows from a distance (fig. 22). compound pahoehoe lava flows. the lava pile is built up of compound pahoehoe flows in packages that may reach thicknesses exceeding 100 m without any indication of an eruptive pause. the average package thickness in different profiles ranges from 26 m to 55 m. individual flow units range from flow lobes 20 cm thick (fig. 36) to sheet flows 25 m thick, but most flow units are below 5 m and generally 1–3 m thick. thinning of the flow units is apparent in the distal areas on the southern innerit peninsula where the flow units are about 1–2 m thick. the pahoehoe flow units commonly have basal zones that are c. 10 cm thick with pipe amygdales, which are strongly inclined in the flow direction in the upper part of the zone. the orientation of these pipe amygdales has been used to determine the flow direction of the lavas (larsen 1981b). above this basal zone about half of each flow is massive and the upper half is crumbling and highly amygdaloidal, with amygdale banding parallel to the lava top. this banding was formed by upwards concentration of gas released from the still-flowing liquid lava under the solidifying crust. the flow units may tun tun nun nun d d d d d nun fig. 34 subaerial lava flows of the upper part of the nunavik member of the vaigat formation (nun) overlain by subaerial lava flows of tunuarsuk member of the svartenhuk formation (tun) at narsinganersua/kap cranstown, south-western svartenhuk halvø. the sedimentary kuugaartorfik member is not present here. the succession is repeatedly faulted; the apparent low-angle fault is a wnw–ese-trending fault that cuts obliquely through the exposure. note several crosscutting dykes (d) perpendicular to the orientation of the lava flows. the colour differences seen are partly caused by superficial alteration. the exposed section of the nunavik member is 280 m thick. compare fig. 14. photo: asger ken pedersen. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 44 of 121 be continuous through fully exposed fields of view of 50–100 m or more, reflecting their highly fluid nature as well as a planar and nearly horizontal topography during eruption. in other places the flow units are lensoid and of short lateral extent because they represent cross-sections through pahoehoe flow lobes (fig. 36). massive sheet flows. these sheet flows are mostly 4–16 m thick, with a few reaching 25 m in thickness. they are fine-grained and form erosion-resistant horizons. columnar jointing may be present. a scoriaceous top zone forms approximately one-fifth of the thickness of a single lava flow, whereas a scoriaceous base is missing or strongly reduced. the two types of flows are illustrated in fig. 37. subaqueous volcanic rocks. the subaqueous products comprise hyaloclastites, coarser pillow breccias and pillow lavas (fig. 38). the hyaloclastites are commonly characterised by gigantic foreset-bedding and form deposits with a thickness of up to several hundred metres, similar to hyaloclastites of the vaigat formation on disko and nuussuaq (pedersen et al. 2017, 2018). the foreset-bedding dips from 33° to nearly horizontal. steeper dips have also been recorded. the bedding is brought about by trains of material with greater concentration of coarse clasts of pillow fragments interspersed with layers of finer material with fewer of such fragments. the distal bottomsets have low dips and are composed of relatively fine material, and sediment intercalations may be found. in several places (e.g. along simiuttap kuua) the compound pahoehoe lava flows can be followed into steeply north-dipping pillow beds and pillow breccias with foresets with a vertical height of more than 200 m. no contemporaneous coastal erosion of the hyaloclastite layers has been observed, which suggests a very short period of eruption. tuffaceous and other fine-grained volcaniclastic deposits. tuffaceous and soil beds are rare and are generally too thin and poorly exposed to be mapped. most tuff and soil beds are 5–20 cm in thickness and in a subaerial environment they have generally been baked to a brick-red colour by the overlying lava flow. the beds are characterised by an even thickness over the whole surface of the underlying lava flow. together with the thicker sediments mentioned below, these beds represent the only certain pauses in deposition of the lava beds. they are most frequent in the upper part of the member in the central area, where one horizon every 50 m in the upper part of the member has been recorded. the higher frequency of sedimentary beds between flows suggests longer time intervals between eruptions and decreasing magma production rates towards the top of the member. a few thicker sedimentary beds occur. some of these are brown or black; the most significant tuff bed is 8–15 m thick. in the south-eastern area a 14 m thick sedimentary bed is situated 540 m above the base of the member (profile 5). the bed consists of ellipsoidal basalt pebbles at the base, followed by volcanogenic sandstones and shales with plant fossils, e.g. ginkgoites. on the south coast south of saviit qaqqaat, two 4–7 m thick sedimentary beds occur west of the saviit–qooroq fault zone (figs 17, 35). they consist of volcanogenic sandstone and conglomerate with basalt boulders, possibly slide deposits. the fact that coarse clastic beds of local aspect are present in the area west of the saviit– qooroq fault zone may relate them to early uplift of the fig. 35 conglomerate with large clasts of volcanic rocks in a matrix of reddish volcanogenic sandstone and with a red-oxidised soil layer on top. the overlying rock is a massive picrite lava flow. the sediment is interpreted as a massflow deposit formed in connection with early tectonic movements along faults in the saviit flexure zone. field assistant for scale. south coast of svartenhuk halvø beneath saviit qaqqaat, sample number 181052 in fig. 17. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 45 of 121 fig. 36 overlapping flow lobes in a compound pahoehoe lava flow. note the numerous zeolite-filled vesicles and thorough red-oxidation, strongest at the tops of individual flow lobes. height of picture c. 1 m. cliff at kap cranstown, south-western svartenhuk halvø. photo: asger ken pedersen. fig. 37 lava succession comprising alternating groups of thick, massive sheet flows and compound pahoehoe flows built up of many individual, thin flow lobes. the lower pahoehoe flow group is topped by a dark grey sediment horizon. the overlying sheet flows have red soil horizons between them. dykes cut the flows at approximately right angles. height of section c. 60 m. upper part of nunavik member at kap cranstown, south-western svartenhuk halvø. photo: asger ken pedersen. fig. 38 typical coarse hyaloclastite with unsorted pillow fragments in a matrix of glass grains and rich in zeolites. the rock is an olivine-rich picrite. note the 1 cm glass rind on the largest pillow fragment. length of hammer 32 cm. upper nunavik member, southern innerit. photo: asger ken pedersen. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 46 of 121 block east of the fault. on the south coast at tartuusaq, a 2–3 m thick tuff layer is present in the upper part of the member where it appears to reveal a stratigraphic repetition (fig. 16). in the same section a 15 m thick bed of shale and tuff is situated at the upper boundary of the vaigat formation. a crimson-red bed c. 8 m thick of welded basaltic tuff very near the upper boundary of the member was mapped by larsen & grocott (1991) within an area of c. 10 × 3 km in the upper reaches of the valley that leads into tasiusap imaa (figs 26, 39). marker horizons. the hyaloclastites in the basal or lower part of the member form distinct beds, which can be used as marker horizons. tuffs and sediments are well exposed along the south coast where they are useful for short-distance correlation across faults. some lava flows can be used as markers: 1. in the southern, south-eastern and central eastern areas, a distinct, dark green, massive picrite lava flow or sill 15–30 m thick containing equant olivine phenocrysts up to 1 cm across occurs immediately above the nerutusoq member. 2. on the southern side of peak 1309 m in the central area, and in part of the south-eastern area, a characteristic 70 m thick, dark group of flows with many thin pahoehoe flow lobes occurs 300 m above the hyaloclastite (fig. 32). this horizon is easy to identify from a distance and on the oblique aerial photographs and is shown as βo1 on the geological map 1:100 000 svartenhuk. it can be followed over a distance of 26 km from a location 400 m above the hyaloclastite south of peak 1078 m and northwards to the west side of the valley of siuteqqut kuuat, where it goes into hyaloclastite facies (fig. 27). the dark flows also have a distinctly enriched chemical composition (see the geochemistry chapter), and their palaeomagnetic properties are different from those of the underlying grey flows, suggesting a considerable time lapse between the two parts (riisager et al. 2003, 2004). the topographic slope of the flows was 0.9°n. 3. overlying the dark βo1 flow group is a 80 m thick, grey, zeolite-rich group of picrite lava flows comprising thick flow units with segregation veins. it has an areal distribution similar to the βo1 flow group. 4. some picritic sheet flows with olivine-rich nodules occur about 200 m below the top of the member and some tens of metres below a significant tuff horizon in the west at kap cranstown and in the central area on the south side of usuit kuussuat. 5. the uppermost part of the member is dominated by less olivine-rich basalts with a persistent, nearly aphyric, massive basalt flow (βo2) present in many profiles in the central area (fig. 10, profile 24) and at tartuusaq (fig. 16). eruption sites. in general, picritic dykes are spatially restricted to the nunavik member. they are therefore regarded as feeder dykes for the member, and this retun nun fig. 39 the upper part of the nunavik member (nun) containing a crimson-red bed c. 8 m thick of welded basaltic tuff (white arrows) within grey pahoehoe flows near the boundary to the tunuarsuk member (tun). the boundary is traced with white lines. akuleqqut, southern svartenhuk halvø. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 47 of 121 lationship has been observed in a few places. north of kangiusap imaa, a picrite dyke cutting the hyaloclastites shows a gradual change into pillow facies at an altitude above 900 m. on the south side of peak 1309 m, a dyke changes laterally into a picrite lava flow of the βo1 marker horizon (fig. 40). also in this area (for example at 71°44′n, 54°16–17′w, larsen & grocott 1991) dense swarms of thin, irregular, wsw–ene-trending picrite dykes with 1–2 mm sized olivine phenocrysts appear to be feeder dykes for the picritic hyaloclastite as most of them terminate below the lavas. such swarms of feeder dykes terminating in hyaloclastite heaps are similar to eruption sites mapped and described from nuussuaq (pedersen et al. 2007a,b, 2017). the crimson-red bed of welded basaltic tuff mentioned above, which occurs close to the top of the member in a c. 10 × 3 km area in south-central svartenhuk halvø (fig. 39), must have been produced from a local eruption site within that area (approximately 71°34′n, 54°42′w, fig. 26), probably a fissure that was waterlogged at the time of eruption. volcanic evolution of the nunavik member. at the start of the deposition of the nunavik member a water-filled basin was situated along the cretaceous boundary fault system. its extent towards the north and west is unknown because the relevant rocks are below exposure level, and it is also not clear whether the basin was still marine. on nuussuaq 100 km to the south, a similar marine basin had at this time just been sealed off from the sea and changed into a large lake into which hyaloclastites and lava flows of the ordlingassoq member, the correlative to the nunavik member, were deposited (pedersen et al. 1996, 2017). it is possible that the basin north of this seal was still connected to the sea; this remains to be investigated. when the eruption of the nunavik member began, a topographic high of earlier volcanic products existed in the south and south-west, from where subaerial lavas flowed towards the east, north-east and north and into the water-filled basin along the cretaceous boundary fault system. the stratigraphy, inclination of the foreset beds of the hyaloclastites and flow directions of the lava flows (fig. 26; larsen 1981b) all support this picture. the advancing lava flows filled the basin with prograding hyaloclastite deposits several hundred metres thick, and only after this filling was achieved could the lava flows advance further. in this way a large part of the subaerial flows of the southern area were converted into equivalent hyaloclastite breccias towards the north and east. thus, the thickness of the subaerial lava pile decreases to the north and east. at a later stage, the precambrian basement was covered by both hyaloclastites and subaerial lava flows. hyaloclastites. the thickness of the hyaloclastites increases from 50 m to 120 m from south to north along the east coast of svartenhuk halvø. farther north, the hyaloclastites are more than 400 m thick on firefjeld profile 7nunavik member taseraarsuit ßo1 ßo1 nunavik member svartenhuk formation nerutusoq member peak 1309 fig. 40 nunavik member in central svartenhuk halvø around the taseraarsuit lake (blue oval), viewed towards north-west. the nunavik member is underlain by lava flows and hyaloclastites of the nerutusoq member. the marker horizon βo1 is indicated both in the foreground and in the middle ground. other local markers are traced in black on the original photograph. the white lines indicate the boundary between the nunavik member and the svartenhuk formation. profile 7, peak 1309 m, is indicated with a yellow line in the right foreground. the small red dot on the profile line indicates a feeder dyke for the lowest flows in the βo1 marker horizon. geodetic institute oblique aerial photograph 526fn/3759. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 48 of 121 and 300–400 m in the central and north-eastern basin areas at siuteqqut kuuat and the south-west side of the simiuttap kuua valley. west of firefjeld, however, only 50 m of picritic hyaloclastites are found, indicating that a topographic high existed here. to the north-west, the local large depocentre of the nerutusoq member (fig. 26) blocked for deposition of the picritic hyaloclastites of the lower nunavik member. hyaloclastites overlie the precambrian basement at two stratigraphic levels, which both increase in thickness northwards, where they have filled a depression in the highland south-east of siuteqqut. the foreset-bedding of both hyaloclastite horizons indicates northerly and north-easterly transport directions. in conclusion, while the precambrian basement as a whole acted as a barrier for the early volcanism of the vaigat formation it was either partly submerged below sea level, covered by a lake or cut by river channels when the first picrites reached the area or were emplaced locally. after the first infilling and subsequent formation of subaerial lava flows, submergence recurred from the east and caused the formation of the second hyaloclastite horizon. thus, a general subsidence of the eastern and north-eastern part of the precambrian basement took place during this period. the lava pile. the lava pile appears to have formed a lava shield with a low-angle slope of 1° to the north. a study of the flow thicknesses shows that the thickest pahoehoe compound flows are found in the lower part of the member in the south-eastern and central-eastern areas. the thinner compound flows in the upper part of the member presumably reflect decreasing erupted volumes and possibly also thinning of the lava flows in the distal areas. on innerit, the lava flows are thin, c. 1–2 m, and very fine-grained to aphanitic and thus appear to have been chilled quickly. these features are presumably a consequence of the distal position of the flows far from the extrusion sites in the south. thicker sheet flows are much more abundant to the west. in the tasiusaq area they form 37% of the lava flows, and a similar value is found on qinnivik. elsewhere, they represent less than 5%, which could indicate different source areas. chemistry and chemostratigraphy. the nunavik member is strongly dominated by picrites; basalts are subordinate. the total range in mgo is 7.5–29.8 wt% with an average content of 17.5 wt% mgo. other major elements such as tio2 and feo* are correspondingly low relative to the mgo-poorer svartenhuk formation basalts (see the geochemistry chapter). the nunavik member is generally not crustally contaminated, but a few exceptions occur in the lower part of the succession. representative analyses are given in the geochemistry chapter. some flows have slightly higher tio2 contents than the main part (for similar contents of mgo). these occur interspersed within the succession; it has not been possible to establish a detailed chemical stratigraphy of the member. lava flows, which are geochemically enriched in incompatible trace elements, occur sporadically as single flows or groups of flows. the dark picrite flows in the βo1 marker horizon are distinctly enriched in these elements (see data for sample 251527 in the geochemistry chapter). boundaries. the lower boundary of the nunavik member is placed at the base of greyish picrite lava flows or hyaloclastites overlying the brown basaltic lava flows or hyaloclastites of the nerutusoq or kakilisaat members. in reference profile 7 (peak 1309 m), grey hyaloclastites and pillow lavas rest directly on cretaceous sediments, but the relation of these rocks to the nunavik member is uncertain. the upper boundary is defined by the occurrence of the first evolved basalts of the overlying tunuarsuk member of the svartenhuk formation, which have a distinctive brown weathering colour. in the northern area, the boundary is marked by the first deposits of white to yellow sands of the kuugaartorfik member. in this area, there is a small angular unconformity at the boundary as described in the section on the lower boundary of the svartenhuk fomation. geological age. paleocene (danian to selandian).the age of the nunavik member is constrained to the interval 62.2–60 ma based on its eruption later than magnetic polarity chron c27n (riisager et al. 2003) and the age of the overlying svartenhuk formation. the 40ar/39ar age of 61.08 ± 0.56 ma obtained by chauvet et al. (2019) for a silicic basalt flow is within this age interval, regardless of whether the flow is from the nunavik or the kakilisaat member. correlation. the nunavik member is correlated with the ordlingassoq member of the vaigat formation on ubekendt ejland, nuussuaq and disko, based on its stratigraphic position within the lava pile and similarities in the chemical compositions illustrated in the geochemistry chapter (see also figs 83, 84). both these members are relatively heterogeneous and not quite as depleted in incompatible trace elements as the older anaanaa and naujánguit members of the vaigat formation on disko and nuussuaq (larsen & pedersen 2009, 2017). http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 49 of 121 svartenhuk formation new formation history. on the geological map 71 v.1 syd igdlorssuit (larsen 1983), the svartenhuk formation is shown as the units βf3 to βf6. on the geological map 71 v.1 nord svartenhuk (larsen & grocott 1991), the svartenhuk formation comprises the sedimentary unit t (for tertiary) and the volcanic units β1 to β3, i.e. the entire basalt succession overlying the vaigat formation. the name svartenhuk formation was used in the same sense by larsen & pulvertaft (2000), but the formation was not formally defined. larsen et al. (2016) found that the basalt succession consisted of two parts of, respectively, paleocene and eocene age and restricted the svartenhuk formation to comprise only the paleocene part (fig. 3). here the svartenhuk formation is defined to comprise the paleocene part of the succession, whereas the eocene part is defined as the naqerloq formation in a later chapter. name. after svartenhuk (sigguk), the most westerly point on the peninsula of svartenhuk halvø. synonymy. the informal ‘middle formation’ and a part of the ‘upper formation’ of larsen (1981a) and holm et al. (1993). subdivisions. the svartenhuk formation is divided into four members based on lithological and chemical characters (tio2 contents). these are from base to top: kuugaartorfik member: quartzofeldspathic sands, mudstones, coal seams, volcanogenic sediments and tuffs. the sediments are interspersed with hyaloclastites and subaqueous and probably invasive lava flows of the tunuarsuk member. tunuarsuk member: brown basalts, grey olivine-phyric basalts and subordinate dark greenish grey picrites. the lowest part of the member in the northern area comprises hyaloclastites and subaqueous and invasive lava flows, whereas the main part of the member comprises subaerial lava flows. nuuit member: brown aphyric and plagioclase-phyric basalts, mainly subaerial lava flows. skalø member: mostly light grey to yellowish, plagioclase-phyric basalts, subaerial lava flows. distribution. the svartenhuk formation is found in central, northern and western svartenhuk halvø but has been eroded away in the eastern part of the basin area. the formation is widely distributed north of svartenhuk halvø where it occurs as far north as 72°43′n (figs 5, 6). the formation extends from the coastal islands in the west at 56°w and eastwards onto the precambrian basement areas as far east as at least 51°35′w at 72°n (fig. 5); the extent of the volcanic rocks farther east and north-east beneath the greenland ice sheet is unknown. to the west and north-west, the formation is considered to continue on the shelf, which is supported by a seismic section reaching as close as 10 km from the west coast of svartenhuk halvø (gregersen et al. 2013, fig. 5), and by the composition of basalt samples dredged on the upernavik escarpment (fig. 1; polteau & planke 2008; analytical data are given in the geochemistry chapter). in the sea south of svartenhuk halvø, the formation is interpreted as a more than c. 800 m thick package of west-dipping strata seen in seismic section geus00-50 (skaarup & pulvertaft 2007). some volcanic units on western ubekendt ejland and western nuussuaq are also included in the svartenhuk formation based on their age (see the section on correlation). type area. the south-east coast of the innerit peninsula. this coastal stretch shows the most complete section through the formation, with all four members present in a well-exposed, flat-lying, unfaulted succession (fig. 41). here 100–200 m of sediments and subaqueous volcanic deposits are overlain by c. 1350 m of lava flows sampled in profile 64 paannivik (figs 13, 41). the paannivik profile is the type section for the tunuarsuk and nuuit members, whereas the skalø member is incomplete and the kuugaartorfik member is partly exposed 4 km east of the profile. thickness. the sum of the greatest thicknesses of the four members of the svartenhuk formation in the basinal area is 110 + 850 + 480 + 450 = 1890 m. there is not one single area in which this thickness is attained. the most complete profile (no 64 paannivik) has a thickness of c. 1350 m for the volcanic succession, but the upper boundary has been removed by erosion. distal thicknesses in the northern areas are up to 1000 m, for example at qaqqap qaa (figs 9, 13). farther to the northeast, the preserved thickness is c. 600 m at pannertuup qaqqa and 660 m at salliaruseq, in both places constituted by relatively few, thick, presumably ponded flows. in the eastern and north-eastern areas, an unknown thickness of basalt has been removed by erosion and in the most distal areas only remnants of basalt are left. in the domal area east-south-east of svartenhuk halvø (fig. 5), no volcanic rocks are preserved. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 50 of 121 n un av ik m em be r u m iia rfi k pa an ni vi k in ne ri t sk al ø m em be r n uu it m em be r tu nu ar su k m em be r ku ug aa rt or fik m em be r q eq er ta q pa an ni vi k pr ofi le 6 4 a b hy al oc la st ite fi g. 4 1 th e m os t c om pl et e ex po se d se ct io n th ro ug h th e sv ar te nh uk f or m at io n, r es tin g on th e va ig at f or m at io n an d w ith a ll fo ur m em be rs p re se nt . p ro fil e 64 a a nd b a re s ho w n w ith y el lo w li ne s; th e pr ofi le is d ep ic te d in f ig . 1 3. s ou th -e as te rn in ne ri t p en in su la . t he p aa nn iv ik m ou nt ai n re ac he s 12 88 m a lti tu de . g eo de tic in st itu te o bl iq ue a er ia l p ho to gr ap h 52 6h n /5 63 4. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 51 of 121 lithology. the volcanic rocks are brown and grey aphyric and plagioclase + olivine ± clinopyroxene-phyric, mainly subaerial tholeiitic basalts and subordinate olivine-phyric magnesian basalts and picrites. subaqueous lava flows and hyaloclastites are present in the lower part of the formation in the northern and north-eastern areas. quartzofeldspathic sediments are present in the north-western and north-eastern areas. thin horizons of tuffs and sediments are present throughout the formation. eruption sites. in contrast to the picrite successions of the vaigat formation, for which several eruption sites are known, only very few eruption sites have been identified for the basalts of the svartenhuk formation. this is a feature shared with flood basalt successions worldwide, including those on nuussuaq and disko (pedersen et al. 2017, 2018), presumably caused by the effusive character of the fissure eruptions, which fed the lava flows. the oxidised products of more explosive eruptions from central vents are more easily visible, and some sites of this character have been identified. basalt dykes cutting the older parts of the volcanic successions, with compositions matching those of the svartenhuk formation, can be regarded as feeders for this. such dykes exist, but the data are too sparse to give a reliable impression of the extent of the lava production areas. however, no dykes are known more than 10 km east of the boundary fault, and the large volcanic areas to the north and east were presumably sourced from eruption sites within the basin area on svartenhuk halvø and just outside it close to the boundary fault. chemistry and chemostratigraphy. the volcanic rocks of the svartenhuk formation consist of evolved tholeiitic basalts with 3.5–10 wt% mgo and, in the lowest member, also magnesian basalts and picrites with 10–18 wt% mgo. the rocks have low contents of incompatible elements such as k2o (0.08–0.5 wt%) and low p2o5/tio2 ratios (0.07–0.12). a few scattered lavas enriched in incompatible elements occur throughout the succession. representative analyses are given in the geochemistry chapter. the formation shows subtle chemostratigraphic variations up-section, which are recognisable throughout the area and have been an important help in the validation of the field-based stratigraphy. the variations are best seen in plots of tio2 contents vs. height in sample profiles as shown in fig. 42. boundaries. the lower boundary in southern and central svartenhuk halvø is placed at the first occurrence of subaerial, brown-weathering, olivine-poor, plagioclase-phyric basalts above the main picrite series of the underlying vaigat formation. in northern svartenhuk halvø, where subaqueous and fluvial conditions prevailed, the lower boundary is placed at the base of a succession of white quartzofeldspathic sand, volcanogenic sediments and olivine-poor hylaoclastites that overlie the picrites of the vaigat formation. tectonic movements prior to the emplacement of the svartenhuk formation (larsen & pulvertaft 2000) gave rise to minor unconformities between the two formations. this is seen at two localities labelled h and i, shown on fig. 6. loc. h: between 71°36′n, 54°33′w and 71°35.5′n, 54°29′w along the southern side of the central part of usuit kuussuat, 11–17°sto se-dipping picrite flows underlie 5–8°wto sw-dipping flows of the svartenhuk formation. loc. i: on the west side of the kuugaartorfik mountain east of the inner umiiarfik fjord and north-east of the cretaceous boundary fault system, lava flows of the nunavik member dip 8–10°ne, whereas the svartenhuk formation is nearly horizontal with low nw-dips. an unconformity is seen on south-eastern innerit between locs b and e (fig. 6). here, lavas of the nunavik member appear to dip nw whereas the svartenhuk formation has a low dip of 2–4°sw, decreasing up-section. this is interpreted as a depositional unconformity, with the lavas of the svartenhuk formation drowning the domed lava shield of the vaigat formation. the upper boundary is defined in the sigguk profile in western svartenhuk halvø (fig. 11). the boundary here is placed at the base of a series of brown-weathering, geochemially enriched basalts of the naqerloq formation (arfertuarsuk member) characterised by higher contents of incompatible elements than the basalts of the svartenhuk formation. the brown basalts rest on a poorly exposed tuffaceous sediment horizon 5–10 m in thickness, which constitutes the top of the skalø member of the svartenhuk formation. in south-western svartenhuk halvø north of arfertuarsuk, a similar, but coarser, volcaniclastic sediment horizon is c. 20 m thick. it is overlain by 0–2 flows with chemical compositions corresponding to the naqerloq formation, followed by the thick marker flow of the arfertuarsuk trachyte. in the upper reaches of the qooruusaq valley at c. 180 m a.s.l. the sediment horizon (inferred, 10–15 m unexposed) is directly overlain by an erosional remnant of the arfertuarsuk trachyte. the boundary is also present on skalø; it has not been mapped there and is not shown on the 1:100 000 map of larsen & grocott (1991) but was identified by sampling and chemistry. its inferred position is shown on fig. 6. geological age. paleocene (selandian to thanetian). radiometric age determinations (40ar/39ar) of the svartenhuk formation range from 59.78 ± 0.41 ma http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 52 of 121 vaigat formation naqerloq formation lorem ipsum tunuarsuk member arfertuarsuk member svartenhuk formation 0 1 2 3 4 5 tio 2 trachyte 36b arfertuarsuk 0 1 2 3 4 5 tio 2 62 sigguk 0 1 2 3 4 5 tio 2 59 -60-61 qooruusaq β3.1 0 1 2 3 4 5 tio 2 41b skalø south β3.1 0 1 2 3 4 5 tio 2 58a amitsoq β3.1 0 1 2 3 4 5 tio 2 55 aputituut 0 1 2 3 4 5 tio 2 30 ivissukk. itinnera sw trachyte level ne 0 1 2 3 4 5 tio 2 52 nuuit qaqqaat 0 1 2 3 4 5 tio 2 50 umiiarfik south 0 1 2 3 4 5 tio 2 64 paannivik 0 1 2 3 4 5 tio 2 51 tunuarsuk 0 1 2 3 4 5 tio 2 44c kuugaartorfik 0 1 2 3 4 5 tio 2 47 peak 1430 m 0 1 2 3 4 5 tio 2 47a peak 1625 m 0 1 2 3 4 5 tio 2 65 salliaruseq 0 1 2 3 4 5 tio 2 40 pannertuup qaqqaa 0 1 2 3 4 5 tio 2 38 qaqqap qaa 0 1 2 3 4 5 tio 2 37 nuup qaava 0 1 2 3 4 5 tio 2 42 marrarnaq 0 1 2 3 4 5 tio 2 fladø tukingasoq w e nuuit member nunavik member tunuarsuk member nuuit member skalø member skalø member nuuit member fig. 42 tio₂ (wt%) for profiles through the svartenhuk and naqerloq formations. the profiles in the lower panel are arranged as projected on a sw–ne line through central svartenhuk halvø. the profiles north of 72°n in the upper panel are arranged as projected on a w–e line around 72°15′n. the vertical scale is the number, starting from 1, assigned to each of the successive flows in a profile. the profiles are aligned on member boundaries, as feasible. dot colours and shapes as in the geochemistry diagrams, see legends in fig. 83. dark green dots in the tunuarsuk member are picrites. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 53 of 121 fig. 42 (continued) tio₂ (wt%) for profiles through the svartenhuk and naqerloq formations. the profiles in the lower panel are arranged as projected on a sw–ne line through central svartenhuk halvø. the profiles north of 72°n in the upper panel are arranged as projected on a w–e line around 72°15′n. the vertical scale is the number, starting from 1, assigned to each of the successive flows in a profile. the profiles are aligned on member boundaries, as feasible. dot colours and shapes as in the geochemistry diagrams, see legends in fig. 83. dark green dots in the tunuarsuk member are picrites. vaigat formation naqerloq formation lorem ipsum tunuarsuk member arfertuarsuk member svartenhuk formation 0 1 2 3 4 5 tio 2 trachyte 36b arfertuarsuk 0 1 2 3 4 5 tio 2 62 sigguk 0 1 2 3 4 5 tio 2 59 -60-61 qooruusaq β3.1 0 1 2 3 4 5 tio 2 41b skalø south β3.1 0 1 2 3 4 5 tio 2 58a amitsoq β3.1 0 1 2 3 4 5 tio 2 55 aputituut 0 1 2 3 4 5 tio 2 30 ivissukk. itinnera sw trachyte level ne 0 1 2 3 4 5 tio 2 52 nuuit qaqqaat 0 1 2 3 4 5 tio 2 50 umiiarfik south 0 1 2 3 4 5 tio 2 64 paannivik 0 1 2 3 4 5 tio 2 51 tunuarsuk 0 1 2 3 4 5 tio 2 44c kuugaartorfik 0 1 2 3 4 5 tio 2 47 peak 1430 m 0 1 2 3 4 5 tio 2 47a peak 1625 m 0 1 2 3 4 5 tio 2 65 salliaruseq 0 1 2 3 4 5 tio 2 40 pannertuup qaqqaa 0 1 2 3 4 5 tio 2 38 qaqqap qaa 0 1 2 3 4 5 tio 2 37 nuup qaava 0 1 2 3 4 5 tio 2 42 marrarnaq 0 1 2 3 4 5 tio 2 fladø tukingasoq w e nuuit member nunavik member tunuarsuk member nuuit member skalø member skalø member nuuit member http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 54 of 121 (chauvet et al. 2019), 60.31 ± 1.39 ma and 59.05 ± 0.61 ma for the lowermost part (tunuarsuk member) to 57.98 ± 0.59 ma for the upper part (skalø member) of the formation (larsen et al. 2016). 40ar/39ar age determinations of the maligât formation above the vaigat formation in the disko–nuussuaq area range from 61.2 ± 0.4 ma to 60.2 ± 0.5 ma (storey et al. 1998). this shows that most of the svartenhuk formation is younger than the maligât formation and that there is a possible overlap in time for the oldest basalts of the svartenhuk formation (fig. 3). correlation. the qeqertalik and tuperssuartâta kûa members on ubekendt ejland and the nûluk and ifsorisoq members on nuussuaq west of the itilli fault have been placed within the svartenhuk formation based on their radiometric ages (fig. 3; larsen et al. 2016). the thick tuffaceous sediment horizon at the top of the svartenhuk formation is tentatively correlated with the sediments of the ifsorisok member on nuussuaq (hald 1976; larsen et al. 2016). the horizon is thus potentially of regional extent and is interpreted as representing a break in the volcanic activity. kuugaartorfik member new member history. the member is shown as intravolcanic sediments (t) on the geological map 71 v.1 nord svartenhuk (larsen & grocott 1991). name. after the mountain kuugaartorfik in northern svartenhuk halvø. distribution. the kuugaartorfik member is exposed in northern svartenhuk halvø and the south-eastern part of the innerit peninsula, in both areas resting on lava flows of the vaigat formation (larsen & grocott 1991; fig. 6). north of 72°n, along the coasts of the innerit inlet, sediments of the member are exposed at sea level below the lavas of the svartenhuk formation. isolated outcrops of the member rest on basement gneiss on both sides of sullua, the northern arm of the umiiarfik fjord (fig. 6). the southern delimitation is depositional; to the west the member disappears below exposure level and to the north and east its extent is unknown. type section. kuugaartorfik (profile 44c, fig. 12), located in a well-exposed west-facing gully on the north-west slope of kuugaartorfik mountain between 380 and 490 m a.s.l. (fig. 43). coordinates in appendix 1. reference section. profile 44a, just 600 m south of profile 44c. thickness. the sediment-bearing interval is up to 110 m thick as in the type section, which consists of 80 m of sediment and 30 m of lava flows (fig. 43), but in most areas the accumulated sediment thickness is less than 50 m. lithology. the kuugaartorfik member is composed of white to grey to yellow quartzand quartzofeldspathic sands, black mudstones and coal seams, and grey, ochre or red volcanogenic sand and gravel. the kuugaartorfik member interdigitates with approximately 50% of volcanic rocks including tuffs, lava flows, pillow lavas and hyaloclastites belonging to the tunuarsuk member. the reference profile 44a was visited by both andreasen (1981) and the first author of this volume. it is very similar to profile 44c. the lower part of the member is dominated by 20–50 m of yellow-white, unconsolidated sand layers with minor, thin coal beds overlain by a thick olivine-phyric basalt flow with pronounced columnar jointing. this basalt has a very thin vesicular top zone which, from a distance, makes it look like a sill, but the sediments above are loose with no signs of contact metamorphism. above this lava flow there is a 50 m thick succession of interbedded, unconsolidated quart550 278344 thin basalt flows hyaloclastite entablature columnar basalt columnar basalt sand with thin coal layers olivine-rich picrite ½ m coal 1 m coal 4 m coal with shale 3 m coal with shale 278345 278346 278347 450 500 400 v f ku ug aa rt or fik m em be r tu nu ar su k m em be r m gv04_06_044_lml fig. 43 type section for the kuugaartorfik member in profile 44c, kuugaartorfik. vf: vaigat formation (nunavik member). sediments of the kuugaartorfik member and lava flows of the tunuarsuk member interdigitate. enlarged interval of profile 44c shown in fig. 12. legend in fig. 13. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 55 of 121 zofeldspathic sand horizons, coal seams up to 5 m thick and a single lava flow. the coal seams lack root horizons and contain flat-lying fragments of coal with wood structure, clearly indicating that the coal is allochthonous. coal seams of considerable thickness occur along the south-west side of kuugaartorfik. stratigraphic correlation within the member is difficult due to contemporaneous sedimentation and lava deposition on a flood plain with lakes formed by lava damming, perhaps aided by faulting (andreasen 1981). about 10 km north of kuugaartorfik (loc. j in fig. 6), 130 m of yellow sandstone, siltstone, sand and mudstone beds and a low-grade coal seam cut by a dolerite sill are preserved in a palaeo-valley in the precambrian basement (croxton 1978). a palynological analysis yielded alnipollenites and betulaceiopollenites; the sediments are therefore of paleocene age and were tentatively correlated with the intrabasaltic sediments (here referred to the kuugaartorfik member) by croxton (1978, p. 65). south-west of kuugaartorfik the member is composed of volcanogenic sediments including volcanic pebbles (pulvertaft 1966) with grey, brown to ochre and red colours and coal seams. on the south-west side of a wide valley leading into umiiarfik fjord (loc. k in fig. 6), the member is 40 m thick and rests on an olivine-phyric hyaloclastite of the nunavik member. it comprises a basal coal seam overlain by black tuff and welded tuff (15 m), followed by a brownish volcanogenic sediment (lahaar deposit?) with coal fragments (15 m), and overlain by coal and mudstone scree (10 m). the base of the overlying tunuarsuk member comprises a subaqueous basalt lava flow and hyaloclastite typical of the north-eastern area. the sediments are exposed along the shores of the innerit inlet on the western innerit peninsula around 72°04′n and along the river banks of innerit kuua at the head of the inlet (fig. 44). they have a low westerly dip. the lower part is >20 m thick and comprises cross-bedded yellow sand and gravel layers indicating a fluvial origin with erosion products from the basement. dark siltstones and dark muddy sandstones from near the head of the inlet (samples 456108–456110, fig. 45) contain bisaccate pollen and a few small triporate pollen, all somewhat corroded and difficult to place in age. two corroded examples of the green alga pediastrum, which commonly occurs in lake deposits, support a non-marine or brackish environment (h. nøhr-hansen, personal communication 2020). in the valley c. 4 km east of the bottom of the inlet, on the southern side of the river, the upper part of the member rests on a basaltic sill and is c. 30 m thick. it is composed of grey sand with dark bands rich in garnet (most probably derived from the large proterozoic prøven igneous complex immediately north of innerit), black heterolithic mudstone and a 0.5 m thick coal seam close to the top. a thin whitish layer of devitrified glass is present near the top. this succession is overlain by a thick, subaqueous columnar-jointed basalt lava flow at the base of the tunuarsuk member. the sediments in the valley close to the inlet are disturbed by broad folding, presumably caused by a quaternary glacier, as they are not protected by the lava pile. on the south-eastern coast of the innerit peninsula, the kuugaartorfik member is identified where the top of the picrites is overlain by yellow sand, which is partly concealed by scree of volcanic material (figs 31, 41, 46). the exposed thickness of the member is 50–100 m, but the total thickness must be greater. on the geological map, the top of the member was placed c. 100–300 m above the base in the eastern part of the occurrence. thereby the member occupies a plateau from 300 to 500 m a.s.l. below the steep mountainside with massive basalts of the tunuarsuk member. entablature lavas and hyaloclastites of the lower part of the tunuarsuk member may thus have been included in the ‘intravolcanic sediments’ shown on the 1:100 000 map. the kuugaartorfik member is strongly reduced in thickness or missing in the central part of svartenhuk halvø. the member has a greatly reduced thickness at peak 1180 m on the west side of the nerutusoq valley (pulvertaft 1966; loc. l in fig. 6). a conglomerate with volcanic pebbles occurs south-west of simiuttap kuua. the member has not been identified at the tunuarsuk mountain. sediments possibly belonging to the kuugaartorfik member. a slipped block with at least 10 m of a bedded, coarse volcanogenic sediment with coal clasts is present at sea level by qaqqap qaa. at 190 m a.s.l. above a scree, a similar sediment including coal clasts is referred to the kuugaartorfik member (fig. 13, profile 38). it is overlain by lava sheets, pillow lava and hyaloclastites of the tunuarsuk member. at 300–340 m in the same profile a volcaniclastic breccia interpreted as a massflow deposit is referred to the tunuarsuk member and described later. a number of other sediment horizons occur at or near the boundary between the vaigat and svartenhuk formations. a black sediment horizon (tuff?) 5 m in thickness is present on top of the vaigat formation in the south-western area on eastern aputituut (fig. 10, profile 22) and extends 8 km farther to the south. a black, coarsely fissile sediment and tuff horizon 15 m in thickness is present at the boundary between the nunavik and tunuarsuk members at the south coast west of tartuusaq (fig. 16). origin of the sediments. the sediments of the kuugaartorfik member apparently represent a break in the volcanism and a tectonic event in the northern arhttp://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 56 of 121 akp_1999_01_06svudmod2 nu tun fig. 44 yellow and dark grey sediments of the kuugaartorfik member dipping west, well exposed in the 8 m high coastal cliff on the north side of the innerit inlet on the peninsula of the same name. the lava succession in the mountain wall above the sediments belongs to the tunuarsuk (tun) and nuuit (nu) members. profile 37 (fig. 13) was taken along the shoulder of the wall, which leads up to the 894 m high nuup qaava mountain. photo: asger ken pedersen. fig. 45 siltstone-dominated sediments of the kuugaartorfik member on the north coast of the innerit inlet. the cliff is c. 8 m high. photo: asger ken pedersen. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 57 of 121 eas. presumably the basement to the east was uplifted, and ensuing erosion coupled with flooding led to deposition of quartzofeldspathic sand, mud and coal seams in fluvial plains and lakes in lows in the irregular basement surface and on the exposed surface of the vaigat formation. evidence for tectonic movements after the volcanism of the vaigat formation is provided by the angular unconformity between the vaigat formation and the svartenhuk formation as described here. the mass-flow-like volcanic deposits at qaqqap qaa and south-east of umiiarfik (loc. k in fig. 6) also suggest a period of erosion and significant local topography. the volcanogenic sediments mostly appear to be reworked hyaloclastite material and rounded blocks of basalt, which may have been extruded contemporaneously farther south or west. it appears that the early eruptions of the tunuarsuk member in southern svartenhuk halvø took place simultaneously with the deposition of the kuugaartorfik member in northern svartenhuk halvø, and that the volcanic rocks arrived later in the northern areas. boundaries. the lower boundary is placed at the first occurrence of clastic sediments, including tuffaceous deposits, overlying the main series of picritic lavas of the nunavik member or directly lying on the precambrian basement. the upper boundary is placed at the base of the massive succession of basalts of the tunuarsuk member, although yellow sand sediments similar to those of the kuugaartorfik member occur higher in the sequence with very reduced thicknesses. from a distance, and even close by, it may be diffucult to distinguish between brown soil and palagonitic sediments of the kuugaartorfik member and hyaloclastites of the tunuarsuk member. geological age. paleocene (selandian), close to 60 ma, bracketed between the 40ar/39ar ages of 61.08 ± 0.56 ma for the lower vaigat formation (chauvet et al. 2019) and 60.31 ± 1.39 ma for the lowest tunuarsuk member (larsen et al. 2016). correlation. none known. kuugaartorfik mb coal? nunavik mb tunuarsuk mb fig.sh044_akp_1999_02_09intbassidumimod tunuarsuk mb fig. 46 white quartzofeldspathic, poorly consolidated sandstones of the kuugaartorfik member overlying grey pahoehoe lava flows of the nunavik member of the vaigat formation. the kuugaartorfik member is topped by a black layer of sediment, probably rich in coal, and overlain by lava sheet flows of the tunuarsuk member. south-eastern innerit peninsula beneath the paannivik mountain; compare with fig. 31. photo: asger ken pedersen. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 58 of 121 tunuarsuk member new member history. the tunuarsuk member is shown as the units βf3 and βf4 on the geological map 71 v.1 syd igdlorssuit (larsen 1983). the member is shown as the unit β1 on the geological map 71 v.1 nord svartenhuk (larsen & grocott 1991). name. after the 1260 m high mountain tunuarsuk in the northern part of central svartenhuk halvø. the member is here well exposed from its base on the south-east side of the mountain to its upper boundary where it is overlain by the nuuit member. synonymy. the informal ‘middle formation’ of larsen (1981a) and holm et al. (1993). distribution. the tunuarsuk member is present in a s–nextending area on svartenhuk halvø from the south coast to umiiarfik; it extends northwards on innerit and qeqertaq and eastwards in a widespread but dissected capping on the high precambrian basement (fig. 6). the member marks the onset of extrusion of voluminous lavas that overstepped the cretaceous boundary fault system and spread to the north and east. the member is found as far north as at 72°20′n on qeqertaq (marraarnaq, profile 42) and the inner ukkusissat fjord, but not at pannertuup qaqqaa (profile 40). farther north its presence is undetermined. to the east, recent hand-held aerial photographs (see also fig. 49) and geological maps (guarnieri et al. 2022a, b) show that the tunuarsuk member extends east across ukkusissat fjord to qaarsorsuaq and farther as a series of grey hyaloclastites and lava flows resting on precambrian basement and in some places on sediments of unknown character. the series can be followed east to inngia isbræ but not farther (fig. 5). it has not been visited on the ground. type section. paannivik (profiles 64a and b) on the south coast of innerit (figs 9, 13, 41; coordinates in appendix 1). the lower and middle parts follow the south-western edge of the niaqornaq ridge (profile 64b) from sea level where a thick basalt lava flow with columnar jointing is present, to a height of 575 m a.s.l., while the upper part (profile 64a) follows the edge of the glacier valley north of paannivik from 515 m to 750 m a.s.l.. the base of the member must be just below sea level in the profile itself and is exposed in the cliffs 3.5–9 km north-east of the profile, where the top of the kuugaartorfik member is exposed. reference sections. tunuarsuk (profile 51); this does not include the basal 50–100 m, which, however, are exposed just 2 km farther east. kuugaartorfik (profile 44c), where both the lower and upper boundaries are exposed. aputituut qaqqaat (profile 31) in the southern area, where the member comprises many more thin pahoehoe flows than in the northern areas. profiles in figs 11, 12; coordinates in appendix 1. thickness. a maximum thickness of c. 850 m is found on the south coast of svartenhuk halvø between tartuusaq and illerusat qaqqaat (fig. 16). thicknesses are reduced in central svartenhuk halvø to between 500 m (e.g. at oqaasaq south of tunuarsuk) and 380 m in the mountains between qiterlikassak and usuit kuussuat, possibly because this area formed a topographic high. to the north, the member is c. 700 m thick at tunuarsuk; in southern innerit, it is 750 m thick in profile 64 on the western shoulder of the paannivik mountain (fig. 13) but thins to 300–400 m in south-eastern innerit above the lava shield of the vaigat formation. north of 72°n, the member is c. 500 m thick on western innerit (nuup qaava), >470 m on eastern innerit (qaqqap qaa), and >40 m in northern qeqertaq (marraarnaq). eastwards on the precambrian highland, thicknesses are between 600 and 450 m at kuugaartorfik, siuteqqut and peak 1430 m. farther to the east, the member is around 300 m thick at peak 1625 m, peak 1370 m and peak 1829 m (fig. 7). on the precambrian highland between siuteqqut kuuat and kangiusap aaffaa, where the overlying nuuit member is not preserved, the tunuarsuk member has been at least 300–600 m thick, dependent on the topography and elevation of the basement. lithology. in the following, the subaerial and subaqueous lithologies of the member are described separately. subaerial lithologies. the subaerial lava flows of the tunuarsuk member are mostly brownor grey-weathering, plagioclase-phyric or aphyric basalt lava flows of either compound pahoehoe or sheet flow type, interspersed with dark grey to greenish grey, olivine-rich compound pahoehoe flows of magnesian basalt and subordinate picrite (fig. 47). a few flows appear to be welded spatter flows that presumably occur close to eruption sites. the sheet flows are 5–80 m thick, whereas thin pahoehoe lobes form compound flow groups 10–130 m in thickness. the average lava flow thickness in the individual measured profiles is 9–25 m on svartenhuk halvø, with the thinner flows dominating in the southern and central areas (average 14 m) and the thicker flows dominating in the northern areas (average 22 m). farther to the north (profile 37, nuup qaava) and north-east (profile 38, qaqqap qaa) the average lava flow thickness increases to 33 m and 37 m, respectively, indicating http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 59 of 121 ponding of the lava flows in low-lying areas. sheet flows dominate here (figs 44, 48). the plagioclase-phyric basalts contain glomerocrysts up to 10 mm in size of plagioclase laths intergrown with small olivine grains. the groundmass of the basaltic compound pahoehoe flows is fineto medium-grained and vesicular, with dark green alteration products possibly after glass; from a distance such greenish rocks may be mistaken as olivine-rich types. the basaltic compound pahoehoe flows have plagioclase and olivine phenocrysts that are unevenly distributed between the individual flow lobes; this may indicate magma pulses from stratified magma chambers with very fluid magmas or crystal segregation during flowage, or both. in any case, it indicates a close genetic relationship between the olivine-poor and olivine-rich and picritic magmas. we also consider that the sheet flows and compound flows are closely related genetically. differences are caused by different magma viscosities, commonly in connection with different degrees of evolution. less evolved, mgrich magmas tend to form compound pahoehoe flows, whereas more evolved, mg-poor and ti-rich magmas tend to form massive sheet flows. the lava pile can be roughly divided into three intervals in which lava flows of different morphologies dominate. boundaries between the intervals are gradational. the lower interval (1) comprises a 90–200 m thick succession of brown-weathering sheet flows, mostly of plagioclase-phyric basalt. compound pahoehoe flows are subordinate. particularly in southern svartenhuk halvø, the basal flows of the tunuarsuk member form a group of thick, brown sheet flows, seen for example in the coastal profiles at kap cranstown and tartuusaq (figs 14, 16). to the north, in the umiiarfik region, the lower part is in semior subaqueous facies with two thick entablature lavas constituting the base of the member over a large area. a middle interval (2) forms a 80–370 m thick succession of grey-weathering, interdigitating compound pahoehoe flows and sheet flows. the upper interval (3) is a 190–310 m thick succession dominated by dark greenish-grey compound pahoehoe flows: sheet flows are subordinate. this interval comprises olivine-phyric basalt and a few picrites alternating with grey, aphyric or plagioclase-phyric basalts. in the northern area, a division into a lower succession (interval 1 and 2) of mainly olivine-poor basalts (mgo <8 wt% and tio2 >1.80 wt%) and an upper sucvf tun nuuit tun fig. 47 lava flows of the tunuarsuk member comprising intervals with massive sheet flows interspersed with intervals with thin, crumbling compound flows. height of foreground field of view c. 400 m. in the background, the exposed lava succession comprises the vaigat formation (vf), the tunuarsuk member (tun) and the nuuit member (nuuit) in the 1000–1170 m high mountains. northern svartenhuk halvø looking south-west, with umiiarfik fjord to the right. photo: kristian svennevig. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 60 of 121 cession (interval 3) of mainly olivine-phyric basalts (mgo >8 wt% and tio2 <1.80 wt%) is possible from both field observations and chemical analyses. table 1 gives thicknesses of the lower and upper successions of the member together with the number of lava flows and average thicknesses. whereas thick lava flows of both types dominate in the northern area, lava flows in the southern area are mostly of compound pahoehoe type and generally thinner than in the north. in addition, there seems to be a more frequent interdigitation between mgo-poor and mgo-rich basalts. thus the tunuarsuk member may have been built up by interdigitating flows from two major volcanic sources, one of them with a fluid magma in which transitions from olivine-rich to plagioclase-phyric flow units, or the reverse, can be observed within the same compound lava flow. a southern origin for this suite is suggested. the northern dominance of mgopoor sheet flows in the lower part of the member may reflect a northern origin of this lava type or alternatively that the large erupted volumes filled topographic lows to the north and east. in this case, either the mgo-poor lavas were of the sheet flow type throughout, or degassing of more fluid magma produced the more viscous lava flows. a transition between the two types within one sheet flow has, however, not been seen. remarkably, the highest flow in the member on eastern innerit (at qaqqap qaa) is a picrite with 12.0 wt% mgo that forms a massive sheet flow 50 m in thickness (see fig. 52). subaqueous lithologies. east and north of the cretaceous boundary fault system subaqueous lava flows, pillow lavas, pillow breccias, hyaloclastites and invasive lava sheets are widespread at the base of the member (figs 49, 50). all indicate a general transport direction towards the north. this implies that the whole northern area was more or less flooded before the extrusion of the svartenhuk formation; widespread wetlands existed on the eroded and irregular surface of the nuuit skalø tun fig. 48 the lava succession in the steep east wall of qaqqap qaa, eastern innerit peninsula. profile 38 runs along the ridge to the left through 480 m of light brownish grey tunuarsuk member (tun) flows (left of photo) and 390 m of brown nuuit member flows. light yellowish flows of presumed skalø member were observed from a distance overlying the nuuit member but were not sampled. note the lack of exposures on the lower slopes, where both precambrian basement and sediments of the kuugaartorfik member may be concealed. the water in the lower right-hand corner is the sullua fjord; the peak of qaqqap qaa to the right is at 1210 m altitude. photo: erik vest sørensen. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 61 of 121 precambrian basement and also on top of the vaigat formation. the thickness of the subaqueous deposits increases towards the north, indicating greater water depths here. hyaloclastites and subaqueous lava flows are now exposed from sea level to 1300–1400 m a.s.l. due to uplift and postvolcanic tilting of the lava pile (fig. 5). in the qaqqap qaa profile (38; fig. 13) a coarse volcanogenic sediment with coal clasts at 190–200 m a.s.l. is referred to the kuugaartorfik member. it is overlain by tunuarsuk member pillow lava and hyaloclastites with invasive lava sheets at the base (figs 13, 50). the subaqueous succession is 140 m thick; the upper 40 m is a volcaniclastic, palagonite-rich breccia with transported basalt blocks up to 40 cm in diameter (larsen 1981b, fig. 26), probably a mass-flow deposit (fig. 13). it is here referred to the tunuarsuk member but may as well represent the uppermost part of the kuugaartorfik member. the mass-flow activity may have been caused by movements on a nearby fault in sullua, perhaps connected to the fault in umiiarfik inferred by larsen & pulvertaft (2000, p. 22). in the type profile at paannivik (64) just south of 72°n the basal lava flows have solidified under subaqueous to semi-submerged conditions forming entablature lavas. entablature flows are also seen at the present sea level on kigataq island west of innerit (fig. 51). megapillows up to 4–5 m in diameter occur in the valley east of the innerit inlet. subaqueous deposits do not occur in the central and southern areas. eruption sites. eruption sites for the tunuarsuk member have rarely been observed. welded spatter flows indicating proximity to an eruption site occur in profiles in south-central svartenhuk halvø, e.g. at peak 1050 m (profile 29 at 885–900 m, not shown) and aputituut qaqqaat (profile 31 at 750–760 m, fig. 11). also in profile 31 at 872–890 m a deposit of more or less welded, glass-rich, palagonitised tuff 18 m in thickness forms the uppermost unit in the tunuarsuk member. the deposit has in part the character of a basaltic ignimbrite and is recognised in other profiles in southern svartenhuk halvø. an analysis of the glass (sample 164917 from profile 18 at kussineq) with relatively high mgo (9.1 wt%) and low tio2 (1.9 wt%) confirms the affiliation to the tunuarsuk member. the deposit is probably related to an eruption site in the table 1 tunuarsuk member thicknesses and number of flows with greater and lesser than 8 wt% mgo profile lower part of member upper part of member whole member total member thickness thickness no.of flows no.of flows ave. flow thickness thickness no.of flows no.of flows ave. flow thickness accumulated thickness no. name mgo ≥8.0% mgo <8.0% mgo ≥8.0% mgo <8.0% mgo ≥8.0% 39 peak 1540 m 320 125 0 3 42 195 3 4 28 94 38 qaqqap qaa 470 263 0 5 53 207 3 1 52 154 37 nuup qaava 434 254 0 7 36 180 3 3 30 70 47 peak 1430 m 440 120 1 4 24 320 9 4 25 214 44c kuugaartorfik 570 320 1 12 25 250 9 6 17 140 64 paannivik 810 575 1 24 23 235 10 2 20 170 50 umiiarfik south >520 >330 1 16 19 190 8 4 16 114 51 tunuarsuk 600 325 5 11 20 275 10 4 20 206 30 ivissukkat itin.a 294 134 1 8 15 160 10 4 11 98 29 peak 1050 m 334 158 3 5 20 176 8 5 14 138 31 aputituut qaq. 420 90 0 7 13 330 4 10 24 142 22 akuleqqut 400 150 1 15 9 248 8 7 16 160 25 equuttat k.k b. >330 >157 >3 >9 13 >173 3? 4? 25 108 no.: number. ave.: average. a ivissukkat itinnerat. b equuttat killiit kuuat. all thicknesses in metres. distinction between flows with mgo <8.0% and mgo ≥8.0% is based on a combination of chemical analyses and visual criteria such as phenocrysts, colour and morphology. profiles are arranged from north to south. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 62 of 121 area at 71°31′n, 55°00′w (fig. 26). here a one-kilometre-long, wnw–ese-oriented, 10 m wide, dyke-like row of ochre-coloured palagonite breccias with black glass clasts, presumably near-surface crater deposits, cuts lava flows of the tunuarsuk member. the row is indicated by a line of stars on the geological map of larsen & grocott (1991). the breccia is interpreted as having formed by interaction between a magma erupting through a fissure and ground or surface water. chemistry and chemostratigraphy. the tunuarsuk member comprises tholeiitic basalts with 5–10 wt% mgo, magnesian basalts with 10–12 wt% mgo and scattered picrites with 12–22 wt% mgo. the average mgo content is 8.4 wt%. tio2 is in the range 1.1–3.3 wt% with an average of 1.8 wt%. a few flows scattered throughout the area are distinctly enriched in incompatible elements, and very few flows are crustally contaminated (see the nomenclature section for definition of these types). representative analyses are given in the nuuit la pc sub1 sub2 hy tun fig. 49 volcanic rocks of the svartenhuk formation overlying precambrian metasediments of the karrat group (pc) at around 1000 m altitude. the volcanic succession of the tunuarsuk member (tun) comprises two dark subaqueous lava flows in entablature facies (sub1 and sub2), overlain by light grey hyaloclastites (hy), which presumably displaced the water in a local basin and are overlain by associated subaerial lava flows (la). the overlying brown flows of the nuuit member are all subaerial. the uppermost flows are grey and are assumed to belong to the skalø member, but this part of the succession has not been visited on the ground. there are traces of sediment at the base of the volcanic succession and between the subaqueous flows. western wall of ukkusissat fjord around 72°05′n, 53°40′w; the top level in the background is around 1700 m altitude. photo: erik vest sørensen. vs vs sh047,5_jgl_shfoto_qaqqapqaa fig. 50 irregular, columnar-jointed lava sheets invading volcanogenic sediments (vs). lower part of the tunuarsuk member, base of profile 38, qaqqap qaa, at 190 to c. 220 m altitude (fig. 10). http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 63 of 121 geochemistry chapter. the most complete chemical profiles are those of paannivik (profile 64), umiiarfik south (profile 50), tunuarsuk (profile 51), kuugaartorfik (profile 44c), peak 1430 m (profile 47) and ivissukkat itinnerat (profile 30). there is generally a decreasing tio2 content with height in all profiles (fig. 42), ranging from about 2.5 wt% tio2 (maximum 3.3 wt%) in the lower part to about 1.5 wt% tio2 in the upper part with the more magnesian basalts. to the north, where lava flows with higher tio2 contents are lacking, the variation is best described as a slightly bumpy trend of decreasing tio2 with height. in contrast, profile 30 in the central area exhibits two consecutive intervals each starting with high and ending with low tio2 values, followed by a third interval with almost constant tio2. the different patterns of stratigraphic development in the northern and southern areas may indicate the existence of two separate or poorly connected, deep-seated magma chambers. boundaries. the base of the tunuarsuk member is defined by the first occurrence of brown, olivine-poor, dominantly plagioclase-phyric lava flows or hyaloclastites overlying the picrites of the vaigat formation. in areas where sediments of the kuugaartorfik member are present, the two members interdigitate. in the central area (usuit kuussuat) a single, dark grey, sparsely olivine-phyric to aphyric basalt flow (βo2) interbedded near the top of the picrites has been included in the underlying nunavik member. the tunuarsuk member appears to overlie the nunavik member conformably in most areas, although there are exceptions as described in boundaries of the svartenhuk formation. the upper boundary has been placed at the top of the succession of the olivine-phyric basalts underlying thick, brown basalts of the nuuit member (figs 48, 52). geological age. paleocene (selandian to thanetian). radiometric (40ar/39ar) age determinations of the tunuarsuk member yielded ages of 60.31 ± 1.39 ma (larsen et al. 2016) and 59.78 ± 0.41 ma (chauvet et al. 2019) for the lower part of the member and 59.05 ± 0.61 ma for the middle part of the member (larsen et al. 2016). correlation. the tunuarsuk member is suggested to be fig. 51 lava flow with a basal colonnade of coarse columns along the shore, changing upwards to slender columns in an entablature zone with a wavy column pattern. the overlying part of the flow is thoroughly brecciated. despite these signs of emplacement into water, the flow is capped by a red-oxidised sediment layer (white arrow) indicating that the top of the flow was above water level after emplacement. height of field of view c. 15 m. east side of kigataq island. photo: asger ken pedersen. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 64 of 121 time-equivalent to the qeqertalik member on ubekendt ejland and the nûluk member on nuussuaq west of the itilli fault (larsen et al. 2016; fig. 3). nuuit member new member history. on the geological map 71 v.1 syd igdlorssuit (larsen 1983), the nuuit member is included in the unit βf6. on the geological map 71 v.1 nord svartenhuk (larsen & grocott 1991), the member is partly shown as the green unit β2. in southern svartenhuk halvø where the boundary between the nuuit member and the skalø member (β3) was not mapped, lavas of the nuuit member are included in the blue unit (β3) and labelled β2.1. in the north-eastern area of the map, an upper boundary of the nuuit member was not definable during mapping, and here the nuuit member has been given the blue colour of the skalø member and is not labelled. name. after nuuit, the point forming the northern corner of the inlet amitsup sullua on the north-west coast of svartenhuk halvø. the member is well exposed on the slopes of the nearby mountain nuuit qaqqaat. distribution. the nuuit member has been recognised in southern, central and northern svartenhuk halvø from tasiusap imaa in the south to the north coast and farther north on innerit and qeqertaq and the coastal islands (fig. 6). in north-western svartenhuk halvø (north of maligissap kuua) and south-western skalø, the member is below exposure level. the member is present in south-western svartenhuk halvø (qinnivik) west of the large fault through the arfertuarsuk inlet, and it is also present in southern svartenhuk halvø north of tasiusap tun tun nuiit nuuit fig. 52 the boundary between lava flows of the tunuarsuk member (tun) and the more brownish lava flows of the nuuit member in profile 38 (qaqqap qaa). the thick, columnar-jointed uppermost flow in the tunuarsuk member is mg-rich with 12.0 wt% mgo, i.e. it classifies as a picrite despite its massive morphology. the flow is 50 m thick (fig. 13). the boundary is seen at a distance in fig. 48. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 65 of 121 imaa; however, in these two areas, it has not been mapped because of poor exposures and strong faulting. on the map of larsen & grocott (1991) the nuuit member in these two areas is included in the western part of the skalø member area (β3). in fig. 6, we have drawn a tentative boundary between the nuuit and skalø members in southern svartenhuk halvø, but this was not possible for the qinnivik area. in eastern svartenhuk halvø, the nuuit member is eroded away west of the cretaceous boundary fault system. the member reappears east and north-east of the boundary fault where it overlies the tunuarsuk member (profiles 44c, 47 and 38). north of around 72°15′n (profiles 40 and 65) the nuuit member has overstepped the tunuarsuk member and rests directly on basement gneiss (figs 6, 13). its presence north of 72°20′n is undetermined. in the unvisited areas east of ukkusissat fjord, recent aerial photographs (fig. 49) and geological maps (guarnieri et al. 2022a, b) show that a series of grey flows of presumed tunuarsuk member is overlain by brown lava flows with a visual character similar to flows of the nuuit member to which we presume they belong. these brown flows extend east to inngia isbræ and farther (fig. 5; guarnieri et al. 2022a, b). east of inngia isbræ the nuuit member appears to be the only member of the svartenhuk formation present. a single reconnaissance sample from a nunatak above the headland of the glacier umiammakku sermia at 72°04.2′n, 51°54.5′e (fig. 5, sample 566505) is virtually chemically identical to the uppermost flow in the salliaruseq profile (65) at 72°25.0′n, 53°47.9′e (sample 566511). considering the distance of 77 km between the two localities it is possible, but not provable, that the two represent the same lava flow. type section. paannivik (profile 64a, figs 13, 41) on southern innerit. here the nuuit member is present between 750 m and 1210 m a.s.l., and both base and top of the member are exposed. coordinates in appendix 1. reference sections. nuuit qaqqaat (profile 52) between 120 m and 600 m a.s.l.; the base of the member is not exposed (fig. 53). umiiarfik south (profile 50) between 650 m and 1000 m a.s.l.. this profile is of relatively easy access near the coast (fig. 27), and both base and top of the member are exposed. profiles in fig. 12; coordinates in appendix 1. thickness. the member is 460 m thick in the type section. thicknesses of 300–500 m are commonly found, e.g in the nuuit qaqqaat, umiiarfik south and aputituut profiles (figs 11, 12; table 2). thicknesses in the northern areas are similar or greater, with 460 m at paannivik, >380 m at nuup qaava, >400 m at qaqqap qaa, and >550 m at marraarnaq. in the far north-east the formation has even greater thicknesses, with >580 m at salliaruseq and >630 m at pannertuup qaqqaa; the latter profile comprises thick sheet flows, subaqueous entablature flows and a hyaloclastite bed more than 100 m thick that was fed by a compound pahoehoe flow. the increased thicknesses in the north are probably caused by ponding of flows in more or less water-filled depressions. the lava flows, or at least some of them, came from the south as indicated by the north-dipping foreset-bedding of the hyaloclastites north of qaqqap qaa and at pannertuup qaqqaa. lithology. the nuuit member is dominated by brown-weathering aphyric and plagioclase-phyric basalt flows that generally form relatively thick sheet flows with thick scoriaceous tops with reddish-orange colours (fig. 54). most flows are subaerial but subaqueous units are present, mainly in the northern areas, e.g. hyaloclastites at pannertuup qaqqaa (fig. 13, profile 40). a few tuff layers are also found. in the type profile at paannivik, the lower part of the nuuit member is dominated by sheet flows up to 50 m in thickness of plagioclase-glomerophyric and less frequent aphyric basalt. the lowest flow has entablature structure. higher up there are a few compound pahoehoe flows c. 30 m thick. in the profile at nuuit qaqqaat a pillow horizon is present. in the umiiarfik south profile several tuff beds up to 1 to 2 m thick are present. the northernmost coastal exposures of lava flows table 2 nuuit member thicknesses, number of lava flows and average flow thicknesses profile total no. of average flow no. name thickness lava flows thickness (m) (m) 65 salliaruseq >590 17 35 42 marraarnaq >500 c. 13 38 40 pannertuup qaqqaa >560 c. 13 43 39 peak 1540 m 450 13 35 38 qaqqap qaa 380 11 35 37 nuup qaava >340 15 23 64 paannivik 460 19 24 50 umiiarfik south 350 21 17 52 nuuit qaqqaat >480 29 16 62 sigguk >250 16 16 55 aputituut 330 20 17 31 aputituut qaqqaat 280 19 15 profiles are arranged from north to south http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 66 of 121 are found around 72°22′n on the western tip of the qulassivik peninsula, 20 km east of the kangersuatsiaq settlement (prøven, fig. 7). here lava flows of the nuuit member ran into water and developed megapillows up to 5 m in diameter, spectacular colonnades, tiered colonnades (figs 55, 56) and entablatures before they were banked up against steep hillocks of the precambrian basement (fig. 57). some horizons form the following markers: • two lava flows in the lower part of the member contain large plagioclase glomerocrysts with stellate outlines as well as individual platy plagioclase crystals c. 1 cm long or larger (profiles 38, 39, 47, 64). • two thick lava flows contain c. 0.5 cm plagioclase glomerocrysts together with platy olivine phenocrysts; these occur at paannivik, on skalø and on the south side of uiarsaariaq (profiles 64, 41b and 54). eruption sites. the average lava flow thickness in individual profiles south of 72°n varies between 15 and 24 m, with the lower values found in southern and central svartenhuk halvø and the larger values on the innerit peninsula. average lava flow thicknesses north of 72°n range from 23 to 43 m (table 2). this situation is similar to that in the tunuarsuk member of the same profiles, suggesting that the eruption sites for both members were situated close to or within the southern areas, giving rise to many small and thin flows in the south, whereas only the most voluminous flows reached north to the innerit area and farther to the northern and north-eastern areas. a volcanic plug was found by münther (1973) in the coastal cliffs on the western side of the arfertuarsuk inlet in south-western svartenhuk halvø (figs 26, 58). the plug cuts vertically through lava flows of the tunuarsuk member and one or two flows of the nuuit member. it has a doleritic core and a chilled margin with weak columnar jointing perpendicular to the outer contacts. towards the highest flow in contact with the plug, a purple to blackish agglomerate is developed that passes skalø skalø member nuuit qaqqaat profile 52 profile shift innerit nuuit member fig. 53 the nuuit and skalø members of the svartenhuk formation at nuuit qaqqaat, looking nnw. the white line is the boundary between the two members. some local marker flows are traced in black dotted lines on the original photograph. profile 52 (fig. 12) is indicated with a yellow line. height of section 780 m. geodetic institute oblique aerial photograph 526hn/5641. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 67 of 121 laterally into a lava flow, one of the lowest flows in the nuuit member. higher lava flows are unaffected. in the scree below the plug some gabbro nodules were found. further description is given by münther (1973, p. 15). chemistry and chemostratigraphy. the nuuit member comprises tholeiitic basalts with 3.4–11 wt% mgo; the average mgo content is 6.4 wt%. tio2 is in the range 1.5–3.8 wt% with an average of 2.4 wt%. a few scattered flows are crustally contaminated. the magmas were relatively depleted in incompatible trace elements such as rb, ba, th, u, nb, and light ree. some flows in the northern areas are enriched in the same elements, which has stratigraphic significance as discussed in the later section on petrology. representative analyses are given in the geochemistry chapter. a common pattern of tio2 variation with height in the nuuit member is seen in several profiles (fig. 42). with some irregularities, the lower part of the member shows constant to increasing tio2 with height, reaching a maximum around the middle of the member. tio2 decreases in the upper part of the member, until a reversal to increasing tio2 is seen in a few flows at the top of the member. this is particularly clear in the nuuit qaqqaat and paannivik profiles. if this is a general trend, the upper part with decreasing tio2 seems to be missing in the sigguk profile, whereas the reversal at the top is well developed (the lower part of the nuuit member is not exposed in this profile). similarly, the erosional remains of the member in the eastern profiles kuugaartorfik and peak 1430 m represent the lowest part of the member, whereas the northern profile at qaqqap qaa seems to represent a condensed succession. the umiiarfik south profile is different, with two cycles of decreasing tio2 present. the basalts in the aputituut profile are characterised by relatively high tio2 contents of 2.5 to >3.0 wt% in both the nuuit member and the overlying skalø member (fig. 42). the boundary between them is placed where the flow morphology changes to flows with yellow-brown top zones and plagioclase aggregates typical of skalø member flows. boundaries. the lower boundary is placed where the dark greenish-grey olivine-phyric basalts of the tunuarsuk member give way to brown sheet flows of plagioclase-phyric and aphyric basalt. the boundary is well exposed and easy to identify in most areas due to the colour contrast between the two members. the upper boundary is defined by the appearance of grey-weathering, thin pahoehoe lava flows with yellow flow tops of the skalø member. the boundary is well exposed in north-western svartenhuk halvø around the estuary and river of amitsup sullua and amitsup kuua nuuit nuuit tun tun umiiarfik qooroq fig. 54 characteristic brown, massive lava flows of the nuuit member overlying thinner and more variable lava flows of the tunuarsuk member (tun). northern svartenhuk halvø with the qooroq valley to the left, looking north towards the umiiarfik fjord. photo: asger ken pedersen. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 68 of 121 fig. 55 thick lava flow with a 20–25 m high colonnade with bent column tops. nuuit member at asungasungaa, west coast of qulassivik peninsula, 9 km east of the kangersuatsiaq (prøven) settlement. photo: asger ken pedersen. fig. 56 multi-tiered columns in a thick, ponded lava flow, probably the same flow as in fig. 55. height of section 20–30 m. nuuit member at asungasungaa, west coast of qulassivik peninsula, 9 km east of the kangersuatsiaq (prøven) settlement. photo: asger ken pedersen. pc la la la fig. 57 three successive lava flows (la) of the nuuit member banked up against a steep hill of precambrian basement gneiss (pc). southwest coast of qulassivik peninsula, 10 km ese of the kangersuatsiaq/prøven settlement. photo: asger ken pedersen. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 69 of 121 (fig. 59) and on the south coast of skalø (fig. 60). the top flow of the nuuit member is a distinct, thick, dark brown sheet flow on skalø and at the coast west of amitsoq; this is overlain by a thin sheet flow and yellow-topped pahoehoe flows referred to the overlying skalø member. in the nuuit qaqqaat area, some yellow-weathering lava flows are also present below the top of the member, which is comprised of the same thick, massive, brown lava flow that is the top flow west of amitsoq. on northern skalø, the top of the nuuit member has been placed on top of 2–3 thick brown lava flows overlying some yellow-weathering flows and overlain by thin yellow flows referred to the overlying skalø member. the upper boundary has not been mapped in southern svartenhuk halvø due to poor exposures and many faults. in this area, the geological map of larsen & grocott (1991) shows lavas of the nuuit member labelled β2.1 and included in the blue areas of the overlying younger lavas (β3) with no boundary drawn between them. geological age. paleocene (thanetian). a radiometric (40ar/39ar) age determination of the nuuit member yielded an age of 58.08 ± 0.59 ma (larsen et al. 2016). correlation. the nuuit member is considered to be time-equivalent to the upper part of the qeqertalik member and/or the tuperssuartâta kûa member on ubekendt ejland, and to part of the ifsorisok member on nuussuaq west of the itilli fault (fig. 3; larsen et al. 2016). skalø member new member history. the skalø member is included in the unit βf6 on the geological map 71 v.1 syd igdlorssuit (larsen 1983). it is included in the unit β3 on the geological map 71 v.1 nord svartenhuk (larsen & grocott 1991). name. after the island skalø north-west of svartenhuk halvø where the member is well exposed along the south-east and north-west coasts. distribution. the skalø member has been recognised in north-western and northern svartenhuk halvø, skalø and southern and eastern innerit. it is inferred to be present across all of western svartenhuk halvø, but field relations to the south-west are poorly constrained due tun tun nuuit nuuit fig. 58 volcanic plug feeding one of the lowest lava flows in the nuuit member near the boundary to the tunuarsuk member (tun). the plug is traced in white; the dashed lines are approximate. the central plug is light brown with a diameter of c. 150 m. the upper zone is a purple agglomerate that passes laterally (to the left) into a lava flow. the character of the lower zone is uncertain; it was described by münther (1973) as a “mixed contact”. see text for further description. coastal cliff on the western side of the arfertuarsuk inlet, south-western svartenhuk halvø. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 70 of 121 to intense faulting and poor exposure. the northern and eastern delimitations are erosional. recent aerial photographs (fig. 49) and geological maps (guarnieri et al. 2022a, b) show remnants of a series of lava flows with a visual character similar to flows of the skalø member capping the highest peaks in the area east of sullua and extending east across ukkusissat fjord to inngia isbræ but not farther (fig. 5). these flows rest on brown flows presumed to belong to the nuuit member. they have not been sampled, but their stratigraphic position and comparison with innerit (paannivik and qaqqap qaa) suggest that they belong to the skalø member. as the boundaries to the underand overlying members have not always been mapped, the areas that represent the skalø member on the geological maps (larsen 1983, βf6; larsen & grocott 1991, β3) include parts of the nuuit member in southern svartenhuk halvø and parts of the arfertuarsuk member in western and south-western svartenhuk halvø and on skalø. type section. sigguk (profile 62) on the southern side of the 643 m hill in westernmost svartenhuk halvø; the base is at 150 m a.s.l. and the top is at 360 m a.s.l. (fig. 11; coordinates in appendix 1). both base and top of the member are well exposed, see fig. 64. no other profiles include a well-exposed top of the member. reference sections. south-eastern skalø (profile 41b, fig. 60) contains a marker horizon of many thin pahoehoe units (mapped as β3.1), which is absent in the type profile at sigguk. nuuit qaqqaat (profile 52, fig. 53) contains a 300 m long, well-exposed and well-analysed section through the member. profiles in fig. 12; coordinates in appendix 1. thickness. the greatest thickness encountered is 460 m, found on skalø. in the type section at sigguk, the skalø member is 210 m thick. in central and northern svartenhuk halvø and innerit, erosion has left thicknesses from 300 m at nuuit qaqqaat to 200 m at aputituut qaqqaat to 125 m at umiiarfik south and 130 m at paannivik. thicknesses are summarised in table 3. lithology. the dominant lithology of the skalø member is plagioclase-glomerophyric basalt lava flows of either compound pahoehoe or sheet flow type. the phenuuit member amitsup sullua umiiarfik skalø member fig. 59 the upper lava flows of the nuuit member and the conformable boundary to the skalø member at amitsoq just west of the amitsup sullua inlet, looking se. the water in the foreground is the umiiarfik fjord. note that the lava package west of amitsup sullua is tilted 14°w, whereas the lava package east of amitsup sullua is nearly flat-lying (figs 53, 54). photo: kristian svennevig. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 71 of 121 nocryst contents (commonly >10%) are generally higher than in the other members. the glomerocrysts are c. 5 mm in size or larger. a distinctive basalt type contains characteristic platy olivine phenocrysts and commonly also plagioclase aggregates. this type is found in the middle of the member and is present on skalø and on both sides of the westernmost nw–se-trending fault on the coast between amitsoq and serfat. a similar flow is present on the south-eastern side of the arfertuarsuk inlet (fig. 15, sample 165215). weathering colours are light grey, and a distinct pale yellow colouring of flow tops occurs in many areas, typically on south-eastern skalø and east of amitsup sullua (fig. 61). the yellow colour appears to be due to altered glassy flow tops of relatively thin lava flows or flow units. red flow tops are common in other areas, e.g. in profile 50 (umiiarfik south). the lavas are more grey on south-western skalø. tuff beds are few, and one or two pillow breccias and agglomerates are present on skalø (profile 41a) and at nuuit qaqqaat (profile 52). a single coal seam 0.5 m thick has been recognised on south-eastern skalø (at 310 m a.s.l. east of profile 41b). nuuit mbnuuit mb nuuit mb profile 41b profile shift skalø mb skalø mb β3.1 41b m skalø mb nuuit mbnuuit mb nuuit mb profile 41b profile shift skalø mb skalø mb β3.1 41b m skalø mb fig. 60 light grey lava flows of the skalø member overlying massive brown flows of the nuuit member on the south-east coast of skalø. the succession is faulted (dash-dot lines) and dips 10°sw. the location of profile 41b is indicated (yellow line). marker horizons β3.1 and the dark flow m are described in the text. the coastal section in view is 4.5 km long and the cliffs are 400–450 m high. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 72 of 121 the skalø member contains two marker horizons, which make correlation possible between skalø and the coast of svartenhuk halvø west of amitsoq and one horizon that correlates between skalø (profile 41b) and sigguk (profile 62). a marker flow named β3.1 on the map occurs c. 160 m above the base of the member on skalø (profile 41b, figs 12, 60). it is a prominent, grey, compound pahoehoe lava flow that is almost aphyric in its lower flow units, whereas the upper flow units have large stellate plagioclase aggregates (‘star basalt’) and a relatively coarse groundmass. this unit is 70 m thick on south-eastern skalø and 20 m thick at amitsoq 5 km to the south-east. the flow has higher tio2 content (>3 wt%) than most other skalø member flows and is clearly identified in fig. 42. two ‘star basalt’ sheet flows with a combined thickness of 45 m in qooruusaq (profile 59) have a similar composition and are considered to be part of the same horizon (fig. 42). distinctive, dark flows with platy olivine phenocrysts and commonly plagioclase aggregates occur at two levels in the member on skalø, one above and one below the β3.1 marker (fig. 60; the upper dark flow is labelled m) and along the coast between amitsoq and serfat. the uppermost flow of the skalø member in the two profiles at sigguk (62) and skalø (41b) are compositionally identical and unique for the whole formation (fig. 42). they have anomalously low tio2 contents (1.15 wt%) at 7.6 wt% mgo (mg-number 58.7; see the tio2 diagram in the geochemistry chapter); al2o3 (15.4 wt%) and cao (12.6 wt%) are high but not anomalously so and do not indicate plagioclase accumulation. we consider that these samples, located 19 km apart, represent the same lava flow. in the aputituut area where the boundary between the nuuit and skalø members is difficult to trace, the uppermost c. 70 m of the flows in the aputituut profile (55) have yellow-brown top zones and contain large stellate aggregates of plagioclase phenocrysts; they have therefore been placed in the skalø member. they constitute a coherent group of lava flows with high tio2 contents (2.5–2.9 wt%, fig. 42) and are very similar to the flows in the skalø member marker horizon β3.1. the distance to the latter exposures is c. 25 km, and it is uncertain whether they represent the same horizon. in the same difficult area in central-western svartenhuk halvø, the yellow-weathering colours of the basalts along the river akuleqqut kuuat, and their chemical compositions, confirm the presence of the skalø member (jørgensen 1981). along the north-eastern side of the arfertuarsuk inlet, the sw-dipping succession of lava flows consists of dark grey, olivine-bearing flows and yellow-brown pahoehoe flows with a physical appearance and chemical compositions corresponding to the skalø member, as indicated by three analysed samples (165217–219; fig. 15). because of the presence of olivine in some of these basalts, a small part of the succession was classified as picrite and olivine basalt (βo) on the map of larsen (1983); these flows are, however, normal svartenhuk formation basalts with 6–7.5 wt% mgo. the succession dips below sea level along the eastern side of arfertuarsuk, where the lavas are overlain by sediments some metres thick, followed by a number of brown, swdipping basalt flows that are well exposed in a se-facing cliff (fig.15). these brown basalts also appear to belong to the skalø member according to chemical analyses by agranier et al. (2019, their loc. 11). there is a positive correlation between the total thickness of the member at different localities and the average lava flow thickness (table 3). thus the larger flow thicknesses measured on skalø, at amitsoq and nuuit qaqqaat, and perhaps on southern innerit, may indicate a topographically low-lying area where lava flows could accumulate, in contrast to the environment at sigguk (profile 62) where the individual lava flows are thin, as is the whole member. a horizon of volcaniclastic sediments 5–20 m in thickness is widespread at the top of the member. in the area north of the arfertuarsuk inlet, the sediments are c. 20 m thick (fig. 62) and include rounded cobbles up to 10 cm in size of different basalt types and hyalotuffs. the age of the sediments is not known; the horizon is tentatively included in the skalø member. eruption sites. none have been found. chemistry and chemostratigraphy. the skalø member comprises basalts with 4.5–10 wt% mgo, and the avertable 3 skalø member thicknesses, number of lava flows and average flow thicknesses profile total no. of average flow no. name thickness lava flows thickness (m) (m) 64 paannivik >140 9 15 50 umiiarfik south >120 9 12 52 nuuit qaqqaat >300 17 17 41b skalø south 460 19 24 58 amitsoq >375 19 19 62 sigguk 210 18 12 31 aputituut qaqqaat >200 11 18 http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 73 of 121 age mgo content is 6.5 wt%. tio2 is in the range 1.1–3.5 wt%, and the average tio2 content is 2.0 wt%. the magmas were relatively depleted in incompatible trace elements such as rb, ba, th, u, nb, and light ree, as in basalts of the tunuarsuk and nuuit members. a few flows may be crustally contaminated. representative analyses are given in the geochemistry chapter. the tio2 stratigraphy shows an increase from low tio2 contents of around 1.5 wt% at the base to 2.4% upwards in the member (fig. 42), with the exception of the ‘star basalt’ marker flow (β3.1) in the skalø profile with values above 3% tio2, and a very-low-tio2 basalt with 1.1% tio2 at the top of the member in the skalø and sigguk profiles. the sigguk profile lacks the high-ti β3.1 marker flow and, except for the uppermost low-ti flow, shows a very regular increase in tio2 with height, indicating that the flows form a coherent sequence and presumably were erupted from the same magma chamber. boundaries. the lower boundary of the skalø member is placed at the top of the brown-weathering basalts of the nuuit member with its red flow tops and tuffs. the boundary is distinct in the north-western and northern areas (figs 48, 59–61). it is less distinct and not mapped in the faulted and poorly exposed parts of southern svartenhuk halvø. here, much, or all, of the nuuit member was included in the skalø member area on the map and labelled β2.1, as described previously for the nuuit member. the upper boundary is placed at the base of the brown basalts of the arfertuarsuk member of the naqerloq formation. in the type section at sigguk, a poorly exposed volcaniclastic sediment horizon 5–10 m in thickness constitutes the top of the skalø member. in the area north of arfertuarsuk, the top of the member comprises a volcaniclastic sediment horizon of reworked hyaloclastite material c. 20 m in thickness (fig. 62). in the southern qooruusaq valley a scree-covered horizon 10– 15 m in thickness, presumably of sediment, forms the top of the skalø member and is directly overlain by the arfertuarsuk trachyte flow of the arfertuarsuk member of the naqerloq formation. the sediment horizon probably represents a break in the volcanic activity and is included in the svartenhuk formation. the upper boundary has only been mapped in westernmost svartenhuk halvø in the area between sigguk, svartenhavn and milloorfik. its extent shown on the map of larsen & grocott (1991) is based on long-distance visual correlation and on aerial photographs from the well-exposed boundaries in profile 62 and siggup qoorua (7.5 km north of profile 62, figs 7, 8) and eastwards to the qooruusaq valley. the boundary is also present, though unmapped, in the qooruusaq valley and on skalø. on skalø, it is poorly exposed on the top of the plateau, from where a few analysed samples belong to the overlying arfertuarsuk member. fig. 6 shows our present best estimate of the position of the boundary between the skalø and arfertuarsuk members. geological age. paleocene (thanetian). a radiometric (40ar/39ar) age determination of the skalø member yieldskalø nuuit fig. 61 lava flows of the skalø member with characteristic yellow-tinted, light top zones, overlying brown flows of the nuuit member. north-east side of the amitsup kuua valley, north-western svartenhuk halvø, looking south. the succession is unfaulted and nearly flat-lying, dipping 3°sw. length of cliff section 2.5 km and height of section c. 300–400 m. photo: karl aage jørgensen. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 74 of 121 ed an age of 57.98 ± 0.59 ma (larsen et al. 2016). correlation. the skalø member is considered to be time-equivalent to the tuperssuartâta kûa member on ubekendt ejland and to part of the ifsorisoq member on nuussuaq west of the itilli fault (fig. 3; larsen et al. 2016). fig. 62 heterogeneous volcaniclastic sediments that form the uppermost horizon of the skalø member. height of field of view 14 m; the full thickness of the sediment horizon here is about 15 m. riverbank c. 5 km north of the end of the arfertuarsuk inlet. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 75 of 121 naqerloq formation new formation history. the formation was introduced, but not formalised, by larsen et al. (2016) and comprises those parts of the original maligât and svartenhuk formations, which had at that time turned out to be of eocene age. the reduced maligât and svartenhuk formations, of paleocene age only, are formally revised by pedersen et al. (2018, maligât formation) and in this work (svartenhuk formation). the existing names for the members on nuussuaq and hareøen (kanísut member, hald 1976) and ubekendt ejland (nûk takisôq member, larsen 1977) are unchanged. on svartenhuk halvø, larsen & larsen (2010) introduced two new informal members, the sigguk and arfertuarsuk members, for the eocene part of the volcanic succession. however, new field and compositional data described below have led to new interpretations, and distinction between these two members cannot be upheld. consequently, in this work the naqerloq formation on svartenhuk halvø is not subdivided and the arfertuarsuk member is defined to comprise the entire eocene succession. name. after the 7–8 km long naqerloq valley in westernmost nuussuaq. subdivisions. the naqerloq formation contains three members, which are considered to be approximately laterally equivalent: arfertuarsuk member on svartenhuk halvø and skalø. nûk takisôq member on ubekendt ejland. kanísut member on nuussuaq and hareøen. distribution. the naqerloq formation is present from skalø in the north and southwards across western svartenhuk halvø, through western ubekendt ejland to nuussuaq west of the itilli fault and to hareøen in the south, a n–s distance of around 170 km. the e–w width of the strip of land where the formation is exposed is only 5–15 km. the formation is faulted and variably tilted in all subareas. it is delimited towards the sea or towards faults, and in the north, the eastern delimitation is erosional. type area. the type area is the naqerloq valley in westernmost nuussuaq. the valley is oriented se–nw, it is c. 7 km long and opens into the sea in the northwest. here, a succession of w-dipping lava flows of the naqerloq formation (kanísut member) is relatively easily accessible (hald 1976). reference sections. on ubekendt ejland a reference section in the nûk takisôq member is situated along the south-west coast of the island where the succession dips 32°wsw; the base is situated 5.8 km south-east of the point eqqua (old spelling erqua), up-section is towards nw, and the top of the member is situated 3.0 km south-east of eqqua where it is overlain by lava flows of the erqua formation (larsen 1977). on svartenhuk halvø, the arfertuarsuk profile (36b) and the sigguk profile (62) are described in the section on the naqerloq formation on svartenhuk halvø. thickness. the naqerloq formation is thickest on western nuussuaq where a succession more than 2000 m thick is preserved (hald 1976). on western ubekendt ejland larsen (1977) reported a thickness of c. 900 m. on svartenhuk halvø 300–400 m of succession are preserved. lithology. the naqerloq formation is dominated by massive, brown, subaerial basalt lava flows of sheet flow type. the basalts are plagioclase-olivine-clinopyroxene-phyric, microphyric or aphyric. intermediate and acid tuffs are present at some levels (hald 1976; larsen 1977; larsen et al. 2016), which has significantly contributed to the accuracy of the radiometric age determinations. chemistry. the most important characteristic that distinguishes the basalts of the naqerloq formation throughout the province from the underlying paleocene basalts is their chemical composition. the naqerloq formation comprises tholeiitic basalts enriched in incompatible elements such as k, p, sr, ba, zr, nb and light ree. petrogenetically important element ratios such as p2o5/ tio2, ba/sr, zr/y, nb/y, nb/zr, nb/la, la/sm and gd/lu are all significantly higher in the naqerloq formation basalts than in the paleocene basalts. boundaries. the lower boundary is rarely observed because of faulting and poor exposures. on nuussuaq, the eocene basalts rest on a thick deposit of volcaniclastic sediments, tuffs and mudstones, which form the ifsorisok member of the svartenhuk formation (hald 1976; larsen et al. 2016). on svartenhuk halvø, the eocene basalts similarly rest on a horizon of volcaniclastic sediments 5–20 m thick, referred to the uppermost http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 76 of 121 svartenhuk formation and perhaps a lateral equivalent of the ifsorisok member. the upper boundary is erosional on skalø, svartenhuk halvø and nuussuaq. on ubekendt ejland, basalts of the nûk takisôq member are concordantly overlain by alkali basalts of the erqua formation (larsen 1977). on hareøen, lava flows of the kanísut member are overlain by sediments of the aumarûtigssâ member (hald 1976). geological age. eocene, earliest ypresian. larsen et al. (2016) reported seven radiometric (40ar/39ar) age determinations of basalts from the naqerloq formation, ranging from 57.25 ± 0.95 ma to 54.03 ± 0.33 ma, generally younging upwards (fig. 3). chauvet et al. (2019) reported two basalt ages within the same interval (fig. 3). dating of the arfertuarsuk trachyte flow on svartenhuk halvø has been problematic because of stratigraphic misfit, but this flow is now considered to be situated very close to the base of the formation with an age of 56–57 ma. the flow is considered in detail below. the kanísut member on nuussuaq is reversely magnetised, and the radiometric ages indicate that the magnetochron is c24r (riisager et al. 2003). correlation. a large part of the offshore volcanic succession penetrated in the delta-1 drill hole has ages (56–50 ma) and chemical compositions similar to the naqerloq formation (nelson et al. 2015; see table 7). naqerloq formation on svartenhuk halvø a large part of the naqerloq formation on svartenhuk halvø is situated in a strongly faulted area in south-western svartenhuk halvø. there is no direct connection to the more northerly part of the formation that is less faulted and situated in nearly flat-lying areas immediately east of sigguk/svartenhuk and on skalø (fig. 6). the unique arfertuarsuk trachyte flow near the base of the succession in the southern area is an excellent marker horizon, but, except for a small erosional remnant, it is not present north of maligissap kuua. interpretation of the stratigraphy of the naqerloq formation has therefore been difficult. larsen & larsen (2010) suggested that the succession in the area north of maligissap kuua is older than the trachyte and that the succession in the southern area is younger, with the arfertuarsuk trachyte flow situated at the base of the younger part, stratigraphically around the middle of the combined succession and thus removed by erosion in the northern areas. this interpretation was based on the relations in the qooruusaq valley 12–14 km east of sigguk and c. 20 km north of arfertuarsuk, where the two parts of the succession almost meet each other. in this downfaulted valley the composite qooruusaq profile (59–61, fig. 11) extends from the coast in milloorfik towards the south-east, uphill and up-section through the nw–sestriking, sw-dipping lava pile. it ends at poor exposures less than 3 km north-west of, and below, a northern outlier of the trachyte flow (fig. 6), leaving only 50–75 m stratigraphy below the trachyte unsampled. the sampled flows represent c. 250 m of stratigraphic thickness and all have the enriched geochemical character of the naqerloq formation, apparently supporting the stratigraphic model with the trachyte flow in the middle of the combined succession. larsen et al. (2016) obtained 40ar/39ar ages for a basalt flow on skalø (sample 251372) of 54.86 ± 0.44 ma and a basalt flow in the arfertuarsuk profile (sample 278596) of 55.91 ± 0.60 ma. the arfertuarsuk trachyte flow itself has been repeatedly dated with strangely scattered results ranging from paleocene, 58.4 ma, for the large anorthoclase phenocrysts to eocene, 55.3 ma, for a laser analysis of groundmass alkali feldspar (table 4). the large anorthoclase phenocrysts are sieve-textured and contain many dark inclusions. an explanation of the results may be that these inclusions contain excess argon that was released uniformly through the successive heating steps; there are fewer inclusions in the groundmass feldspar and the inclusions are avoided in the laser analysis. if this is the case, the laser result is the most reliable and the trachyte flow is of earliest eocene age. the question remains, however, how close to the base of the eocene succession is the trachyte flow situated? the northern outlier of the trachyte in the qooruusaq table 4 ⁴⁰ar/³⁹ar age determinations for the arfertuarsuk trachyte flow sample material plateau age ± 2σ (ma) % 39ar released reference 456207 anorthoclase 58.43 ± 0.31 ª 96 author, unpublished data 1931.1 anorthoclase 57.51 ± 0.24 ª 80 larsen et al. (2016) 1931.1 groundmass feldspar 56.85 ± 0.13 ª 31 author, unpublished data s29-e2 alkali feldspar 55.31 ± 1.09 b laser analysis chauvet et al. (2019) all ages are calculated relative to an age of 28.201 ma for the fish canyon tuff standard. ª incremental step heating result from oregon state university dating laboratory. b inverse isochron age. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 77 of 121 valley was briefly visited in 2019 by k. svennevig (fig. 63). below the trachyte, there is an unexposed horizon 10–15 m in thickness, which may represent a sediment horizon although no sediment was seen. two basalt flows below this horizon and a single basalt flow overlying the trachyte were sampled (profile 67 in fig. 9). the geochemical analyses showed that whereas the flow above the trachyte has the enriched character of the naqerloq formation, the two flows below the trachyte have not; they have the typical character of the paleocene basalts and are considered to belong to the skalø member of the svartenhuk formation. the trachyte flow is therefore the oldest eocene lava flow at this locality. the apparent stratigraphic continuity to the eocene flow succession in profile 59–61 farther north in the qooruusaq valley is therefore lost. we must assume that an unexposed fault crosses the valley between the trachyte outlier and the qooruusaq sample profile (fig. 11), with a relative downthrow of c. 300 m of the succession to the north (fig. 6). in the area north of the arfertuarsuk inlet where the trachyte and the overlying lava flows have been mapped (larsen & grocott 1991), the trachyte flow is underlain by two or three lava flows with hyaloclastite and hyalotuffs resting on c. 20 m of reworked volcaniclastic sediment. this sediment has been referred to the uppermost skalø member of the svartenhuk formation. the flows below the trachyte and above the sediments have the chemical character of the naqerloq formation but have significantly lower tio2 (1.8–2.06 wt%) than the flows above the trachyte (>2.4 wt% tio2; see fig. 42 and the geochemistry chapter). flows with similar low-ti compositions occur at the base of the naqerloq formation at sigguk (profile 62) and in the northern qooruusaq valley (profile 59–61). we therefore consider it most probable that the trachyte flow did not reach the northern areas, and that the top of the low-ti flows marks the level corresponding to the trachyte. this means that the three profiles, sigguk (62), arfertuarsuk (36b) and qooruusaq (59–61), are laterally equivalent. indeed, their basalts are compositionally very similar although no single flows are correlatable and the tio2 developments up-section are different (fig. 42). in view of the probable lateral equivalence of the exposed parts of the naqerloq formation, we prefer to define them within the framework of one member only. we therefore extend the arfertuarsuk member sensu larsen & larsen (2010) to cover the entire eocene fig. 63 the northern outlier of the arfertuarsuk trachyte in the uppermost reaches of the qooruusaq valley (fig. 9, profile 67). the exposed trachyte is traced with a solid white line; it is black due to lichen coverage. the presumed total extent of the trachyte is indicated with a dashed white line. see text for descriptions. note the flat country with almost no exposures in the background; this is part of the low-lying area in central western svartenhuk halvø seen in the digital elevation model (fig. 9), where mapping was very difficult and some geological boundaries are only tentative or not placed at all. photo: kristian svennevig. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 78 of 121 succession on svartenhuk halvø and skalø. the sigguk member sensu larsen & larsen (2010) is discontinued. arfertuarsuk member new member history. on the geological map 71 v.1 nord svartenhuk (larsen & grocott 1991), the arfertuarsuk member is shown in the area between sigguk and qooruusaq as the unit β3.3, and in the area between maligissap kuua and the arfertuarsuk inlet as the unit β4. the trachyte flow is mapped as a separate, unlabelled unit. name. after the arfertuarsuk inlet on south-western svartenhuk halvø. distribution. the arfertuarsuk member is present in the western part of svartenhuk halvø from skalø in the north to the bottom of the arfertuarsuk inlet in the south. the member has been mapped in the strongly faulted and tilted areas north-west of arfertuarsuk and traced on aerial photographs on the sigguk peninsula west of the qooruusaq valley (larsen & grocott 1991). however, the formation is also present and sampled on skalø and in the eastern qooruusaq valley where the lower boundary was not sufficiently established to be shown on the map. thus, the western part of the skalø member area (β3 on the map) includes many more areas with lavas of the arfertuarsuk member than shown. our present best estimate of the distribution of the arfertuarsuk member is shown in fig. 6. type section. arfertuarsuk (profile 36b) north of the bottom of the arfertuarsuk inlet. the profile runs across a ridge and follows the south-west side of a wnw–esetrending fault, changing midway to a wsw direction (fig. 9; coordinates in appendix 1). the lava succession is strongly tilted with dips around 20°sw and locally more, and the profile shown in fig. 11 is composed of several parts. the profile contains a thick, mappable trachyte flow, the arfertuarsuk trachyte, near the base of the arfertuarsuk member. reference section. sigguk (profile 62). the lava pile here is unfaulted and flat-lying with a low dip to the east (figs 11, 64). lava flows of the arfertuarsuk member here concordantly overlie lava flows and a sediment horizon of the skalø member and extend from c. 360 m a.s.l. to the geodetic point at 643 m a.s.l. where a distinctly porphyritic flow occurs at the top of the profile. no trachyte flow is present. skalø member arfertuarsuk member sigguk nuuit member profile 62 fig. 64 lava flows of the paleocene nuuit and skalø members of the svartenhuk formation conformably overlain by lava flows of the eocene arfertuarsuk member of the naqerloq formation. profile 62 is indicated with a yellow line. western svartenhuk halvø with the point sigguk (svartenhuk) looking nw. length of the coast from the foreground to sigguk is 10 km. dash-dot lines are faults. geodetic institute oblique aerial photograph 526gn/6272. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 79 of 121 thickness. the preserved thickness of the arfertuarsuk member is c. 300 m in the sigguk area and more than 350 m in the area north of arfertuarsuk. lithology. basalts of the arfertuarsuk member are dominantly brownish sheet flows or compound pahoehoe flows, which are plagioclase ± olivine ± clinopyroxene glomerophyric or aphyric. a few flows with large (>1 cm) platy plagioclase phenocrysts are present at the top of the sigguk profile. sheet flows dominate in the arfertuarsuk profile. flows of the member are mostly glomerophyric, and one flow has box-shaped plagioclase phenocrysts several centimetres in size with reaction rims. two olivine-plagioclase-phyric basalts are present, one of which is a compound flow. one lava flow with relatively high mgo (11 wt%) has ankaramitic affinity with only olivine and clinopyroxene phenocrysts. aphyric basalts dominate the upper part of the profile. several well-exposed tuff beds are present. a dark doleritic basalt flow is present near the top of the succession. in the area north of the arfertuarsuk inlet, a trachyte flow up to 40 m in thickness is present near the base of the arfertuarsuk member (fig. 6). it is the only one of its kind in the region and is important for the stratigraphy of the member, as discussed in the previous section. it was first described and analysed by nieland (1931), who showed it to be an anorthoclase trachyte. the flow is present for 12 km along the se–nw strike direction (larsen & grocott 1991) and thins towards the north. a small erosional remnant is present 7 km farther north in another fault block in the southernmost qooruusaq valley (fig. 6), but except for this, the trachyte flow is missing north of maligissap kuua. the flow is well exposed on the coast at the northern end of the arfertuarsuk inlet. it is composed of several units, of which the basal unit is a more or less welded and devitrified vitrophyre that may be classified as a rheomorphic ignimbrite (jørgensen 1982). the major part of the trachyte is a more normal, massive flow (fig. 65); it is rich in anorthoclase crystals up to several centimetres in size; these large phenocrysts are dotted with dark inclusions in a sieve-texture. microphenocrysts of pale brown clinopyroxene and magnetite are present, and the medium-grained groundmass is very rich in alkali feldspar. the flow contains up to cobble-sized xenoliths of intermediate igneous rock types (sample 278668.1) and syenite (sample 278669.2; fig. 66). the trachyte outlier in qooruusoq contains several 2–6 cm-sized, irregular, fine-grained, dark grey inclusions of extremely evolved composition (sample 568601.2 in fig. 67; see the geochemistry chapter). a fine-grained, dark grey vein cutting the trachyte (fig. 68) likely consists of similar material. eruption sites. none have been found. fig. 65 the arfertuarsuk trachyte flow exposed on the beach, east side of the northern end of the arfertuarsuk inlet. height of section c. 15 m. photo: karl aage jørgensen. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 80 of 121 chemistry and chemostratigraphy. the arfertuarsuk member comprises tholeiitic basalts with 3.8–8.6 wt% mgo and an average of 5.8 wt% mgo. tio2 is in the range 1.8–4.0 wt%, and the average tio2 content is 3.0 wt%. k2o and p2o5 contents are high, with on average 0.80 wt% k2o and 0.39 wt% p2o5. the arfertuarsuk member basalts are also enriched in other incompatible elements such as sr, ba, zr, nb and light ree, and element ratios such as p2o5/tio2, ba/sr, zr/y, nb/y, nb/zr, nb/la, la/sm and gd/lu are significantly increased relative to basalts of the svartenhuk formation. representative analyses are given in the geochemistry chapter. the trachyte flow contains on average 65.7 wt% sio2, 0.83 wt% tio2, 0.24 wt% mgo, 16.4 wt% al2o3, 6.1 wt% na2o and 5.2 wt% k2o. a group of less evolved basalts with low tio2 (around 2 wt%) and p2o5 (around 0.25 wt%) and with the same enriched geochemical character as the more evolved basalts of the member is clearly distinguished in the chemistry diagrams (see the geochemistry chapter). as explained in the earlier section on the naqerloq formation on svartenhuk halvø, these basalts occur below the trachyte flow in the arfertuarsuk area and are used to define an interpreted trachyte level in the northern areas (sigguk and qooruusaq profiles). according to our interpretation, the analysed sample profiles through the eocene basalts from the northern and the southern areas, mainly the sigguk, qooruusaq and arfertuarsuk profiles (62, 61, 36b), are nearly laterally equivalent (figs 11, 42). they are geochemically similar but do not show any obvious mutual correlation. boundaries. the lower boundary is depositional and is placed at the base of a succession of brown lava flows of geochemically enriched basalts and above a volcaniclastic sediment horizon 5–20 m in thickness included in the skalø member. the upper boundary is erosional. geological age. eocene (earliest ypresian), perhaps reaching back into the uppermost paleocene. two radiometric (40ar/39ar) age determinations of basalts yielded ages of 55.91 ± 0.60 ma and 54.86 ± 0.44 ma (larsen et al. 2016), and one of the arfertuarsuk trachyte flow yielded an age of 55.31 ± 1.09 ma (chauvet et al. 2019). see table 4 and the corresponding text for a discussion of the age of the trachyte. correlation. the arfertuarsuk member was erupted in a similar time interval (57–54 ma) and has the same geochemical character as the nûk takisôq member on ubekendt ejland and the kanísut member on nuussuaq and hareøen, and is thus in a general way correlatable with these. fig. 66 feldspar-rich xenolith, probably cognate, of syenite in the arfertuarsuk trachyte flow. length of steel rod 16 cm. east side of the northern end of the arfertuarsuk inlet. photo: karl aage jørgensen. fig. 67 fine-grained, dark grey inclusion in trachyte from the outlier in qooruusaq, cut into two. the inclusion is surrounded by a light-coloured rim almost exclusively of coarse alkali feldspar; contrary to intuition the dark central part consists of about as much feldspar as the rim; the major difference is the grain size. the analysis (278669.2) represents the dark part. length of match 4.7 cm. photo: erik vest sørensen. fig. 68 fine-grained, dark grey vein cutting the arfertuarsuk trachyte flow. east side of the northern end of the arfertuarsuk inlet. length of hammer 45 cm. photo: karl aage jørgensen. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 81 of 121 intrusions svartenhuk halvø is cut by numerous dykes. the dyke distribution map in fig. 69 from larsen & pulvertaft (2000) shows the distribution and directions of dykes as identified on aerial photographs. the map gives an overall impression of dyke patterns, but as stressed by these authors, dykes are much easier to locate in the areas with crumbling picrites of the vaigat formation than in the areas with robust basalt flows of the svartenhuk and naqerloq formations. the relative sparsity of dykes located in the basalt areas in central and western svartenhuk halvø may therefore not be real. the dykes extend north to skalø and southern innerit and also cut the basement and volcanic rocks north-east of the cretaceous boundary fault system, but they quickly die out to the north and east where reconnaissance work revealed very few dykes. the dyke distribution pattern, 54°30'55°00' 54°00' 55°30' 54°30'55°00' 10 200 30 km cbfs: cretaceous boundary fault system dyke with 'eocene' composition dyke or sill with 'paleocene' composition sill sill sill 54°00' 71°45' 71°30' 71°45' 71°30' 71°15' cbfs cbfs gv01.08-018_lmlfig. 69 distribution and directions of dykes on svartenhuk halvø as identified on aerial photographs, from larsen & pulvertaft (2000, fig. 4). the relative sparsity of dykes in the basalt areas in central and western svartenhuk halvø may not be real, see text for explanation. we have added the locations of analysed basalt dykes and sills of two different compositional groups that are supposed to represent two age groups, as indicated in the legend. analytical data from this study and agranier et al. (2019). http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 82 of 121 including the scarcity of tertiary dykes in the northern and eastern areas, confirms the impression gained from the hyaloclastites and lava flows that eruption sites were centred in the basinal areas, from where the lavas flowed to the north and east over distances of up to 80 km for the most voluminous flows. larsen & pulvertaft (2000) presented a detailed structural analysis of the dyke orientations and concluded that several dyke generations are present. the dykes comprise picrites and feldspar-phyric and aphyric basalts petrographically similar to the lava flows, and their chemical compositions (see the geochemistry chapter) comply with the assumption of larsen & pulvertaft (2000) that many dykes are feeders to the lava flows. accordingly, the dykes most commonly have dips perpendicular to the lava flows, which they cut (e.g. figs 15, 16, 37), and the fault zones are generally not intruded by dykes. feeder dykes and eruption sites for all three members of the vaigat formation have been identified in the field and confirmed chemically. in contrast, eruption sites for the svartenhuk formation have rarely been observed, and none are known for the naqerloq formation. picrite dykes and strongly contaminated dykes similar to the kakilisaat member are only found in the areas of the vaigat formation, whereas basaltic dykes are found all over svartenhuk halvø. as described in the geochemistry chapter, the basaltic dykes form two groups with compositions corresponding to those of respectively the svartenhuk and naqerloq formations. this enables us to distinguish between basalt dykes of presumed paleocene age (svartenhuk formation) and of presumed eocene age (naqerloq formation) and thus to confirm the presence of at least two distinct dyke generations. our analysed dykes with naqerloq formation chemistry are all situated in the svartenhuk formation in western svartenhuk halvø, but agranier et al. (2019) published analyses of samples with naqerloq formation chemistry from areas in the vaigat formation along the entire south coast of svartenhuk halvø from tartuusaq to maniiseqqut (fig. 69). this suggests that the eocene lava succession originally extended much farther eastwards on the continent than the present outcrops, as also found for nuussuaq and disko (larsen et al. 2016; pedersen et al. 2018). unfortunately, dyke compositions are unknown in large areas, but it should be possible to map the true distribution of both the eocene and the paleocene dykes by analysing a larger number of dyke samples distributed over the svartenhuk halvø. additionally, even younger dykes may be identified if they are compostionally different from the older lavas and dykes. the cretaceous boundary fault system has apparently facilitated the intrusion of magmas, as seen from a number of large basalt sills intruded close to the fault at simiutaq, along the siuteqqut kuuat valley, near firefjeld and on itsaku (fig. 6). these sills are compositionally identical to the svartenhuk formation and we therefore consider that they are also of late paleocene age. a small intrusion of magnesian basalt is exposed at sea level on the south coast of svartenhuk halvø c. 3.5 km south of tasiusap imaa (fig. 70). it has a cupola-like structure and is c. 20 m wide and 25 m high. it intrudes lava flows of the tunuarsuk member and has a similar chemical composition (10.8 wt% mgo, sample 164863). it is chilled towards the lavas in a c. 5 m wide zone, which has a distinct layering with decimetre-thick, alternating light and dark layers parallel to the contact. a few alkaline dykes (alkali basalt and camptonite) are found on the south coast of svartenhuk halvø (larsen 1983; l.m. larsen 2006). the ages of these dykes are not known and they could be very young. the youngest dykes identified in the region are c. 34 ma old camptonite and monchiquite dykes from ubekendt ejland (larsen 1981c, 1982; clarke et al. 1983; storey et al. 1998). a b fig. 70 an intrusion of magnesian basalt into lava flows of the tunuarsuk member. a: the entire exposure, showing the cupola-like shape. b: the chilled marginal zone with contact-parallel mineral layering. the person (encircled in a) is standing in the same place in both photos. south coast of svartenhuk halvø c. 3.5 km south of tasiusap imaa; the intrusion is shown with a size of 1 × 0.5 mm on the geological map (larsen 1983). http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 83 of 121 geochemistry and petrology of the volcanic rocks identification of crustal contamination and geochemical enrichment the volcanic rocks in the nuussuaq basin consist predominantly of tholeiitic basalts and picrites such as those generated by melting of common asthenospheric mantle (e.g. mckenzie & bickle 1988). however, as mentioned frequently in this study and defined in the nomenclature section, there are also rocks, which are considered to be contaminated with continental crustal material, as well as geochemically enriched rocks with elevated contents of a number of incompatible elements. identification of the two rock groups is best achieved using trace-element ratios. a diagram showing th/nb vs. nb/zr is particularly useful because the crustal-contamination and enrichment vectors are perpendicular to each other (fig. 71). in this figure crustally contaminated samples generally have th/nb >0.12, whereas most enriched samples have nb/zr >0.10. small degrees of crustal contamination and enrichment are difficult to detect. in the vaigat formation (fig. 71a) the kakilisaat member is clearly crustally contaminated, and the extension of the plot field towards the enriched components indicates that some samples are also enriched. the nerutusoq member is enriched but not crustally contaminated. fig. 71b shows that the svartenhuk formation basalts analysed for trace elements are generally uncontaminated although a number of samples have th/nb c. 0.12 and may perhaps be slightly crustally contaminated. there is a trail of enriched samples towards the enriched component field; one sample (262713) has high th/nb (0.16) and low nb/zr (0.09) and is both contaminated and enriched. the enriched character of the naqerloq formation is clearly seen. vaigat formation the volcanic rocks of the vaigat formation on svartenhuk halvø comprise tholeiitic picrites and subordinate basalts. major element variation diagrams for the vaigat formation are shown in fig. 72, traceelement variation diagrams in figs 73 and 74, and ree and multi-element plots in fig. 75. these figures are designed to be comparable with the similar geochemistry diagrams for the vaigat formation on disko and nuussuaq in pedersen et al. (2017). representative chemical analyses are shown in table 5. nunavik member the nunavik member is described first because it constitutes the bulk and ‘normal’ part of the vaigat formation in the area. the two other members are considered to be derived from magmas similar to those of the nunavik member with some additions of material from the crust and upper mantle. the nunavik member is dominated by picrites. the total mgo range is 6.4–29.8 wt% mgo, and the average 0.01 0.10 1.00 10.00 0.01 0.10 1.00 th /n b nb/zr nunavik mb enriched nunavik mb nerutusoq mb kakilisaat mb basement gneiss and sediments ec crustal contamination trace-element enrichment a vaigat formation 0.01 0.10 1.00 10.00 0.01 0.10 1.00 th /n b nb/zr naqerloq formation svartenhuk formation enriched svartenhuk fm basement gneiss and sediments crustal contamination trace-element enrichment ec b fig. 71 th/nb vs. nb/zr diagram showing discrimination between crustally contaminated and geochemically enriched samples. a: vaigat formation. b: svartenhuk and naqerloq formations. the crustal contamination vector points upwards because basement gneiss and sediments have high th/nb, whereas the enrichment vector points towards enriched components (ec) with high nb/zr. data for basement gneiss and sediments from larsen & pedersen (2009). the ec field represents a number of presumed local enriched components. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 84 of 121 mgo content of 130 analyses is 17.5 wt% mgo, which is probably not far from the parental magma composition (larsen & pedersen 2000, 2009; herzberg & o’hara 1998, 2002; herzberg et al. 2007). the well-defined trends, particularly in the major-element diagrams (fig. 72), are caused by olivine fractionation and accumulation, as already noted by clarke (1970). when the fractionating magma reached 7–8 wt% mgo, plagioclase started to crystallise, which removed cao and al2o3 from the magma. figure 72 shows that most of the magmas of the nunavik member did not reach the stage of plagioclase fractionation, with the βo2 basalt marker flow near the top of the member as an exception (6.4 wt% mgo, 13.6 wt% al2o3, 10.6 wt% cao and 2.9 wt% tio2). a couple of samples with low cao also have high sio2 and are crustally contaminated. relatively scattered k2o contents are due to secondary alteration. with some exceptions, the rocks have a relatively depleted geochemical character, meaning that they have low contents of incompatible elements. the incompatible trace elements (figs 71, 73, 75) show a general increase with increasing fractionation. a few scattered flows are enriched in some elements (rb, ba, sr, zr, nb, ree and u); these flows are not distinguishable by their major elements but are clearly distinguished in the trace-element diagrams (figs 71, 73, 75). the dark marker horizon βo1 in the nunavik member has a nb/zr ratio of 0.26–0.28 and is one of the most enriched flows of all (fig. 71); the two samples analysed from βo1 are situated c. 15 km apart. the only trace elements that are fractionated out of the picrite magmas are ni and co in olivine and cr in chromite (fig. 74). the low contents of incompatible elements in the common rocks of the nunavik member are reflected in the ree and multi-element patterns in fig. 75. most ree patterns for the nunavik member have chondrite-normalised (with subscript n) lan <smn so that the la–sm limbs have positive slopes, giving the patterns an overall humpback-like shape. likewise, the multi-element patterns have rb–la limbs with positive slopes. however, some samples have lan ≥smn and near-horizontal to negative slopes of the la–sm and rb–la limbs. the deep troughs at k and pb are characteristic of asthenosphere-derived magmas from the north atlantic tertiary igneous province (e.g. saunders et al. 1997). the distribution of lan/smn ratios of the nunavik member, with most samples having values less than 1 and some samples having values above 1 (see also fig. 83), is similar to that found in the ordlingassoq member of the vaigat formation on nuussuaq and disko (larsen & pedersen 2009; pedersen et al. 2017; see also contour line for the ordlingassoq member in fig. 83). in contrast, the earlier members of the vaigat formation (the anaanaa and naujánguit members) have lan/smn consistently less than 1 and are thus somewhat more geochemically depleted than the ordlingassoq and nunavik members. this is a key feature in the correlation of the last-named members. nerutusoq member the nerutusoq member consists of basalts with 8–12 wt% mgo, normal sio2 contents of 48–49 wt% and relatively high contents of cao (12–13.6 wt%; fig. 72). the rocks have distinctly high p2o5 (0.20–0.32 wt%) and are also enriched in incompatible elements such as ba, th, u, nb, light ree, pb and sr relative to the nunavik member rocks (figs 72, 73, 75). the patterns of enrichment in the nerutusoq member lack the high nb–ta peak and the deep pb trough seen in the enriched rocks of the nunavik member, suggesting that the sources of the enrichment for the two members are different. our analysed samples from the nerutusoq member are not crustally contaminated; however, agranier et al. (2019) reported contaminated, enriched samples presumably from this member (their site 40). it is thus possible that small degrees of crustal contamination have contributed to the differences in enrichment pattern between the two members. the nerutusoq member shows notable similarities to some enriched rock units on nuussuaq and disko (larsen et al. 2003; larsen & pedersen 2009; pedersen et al. 2017). kakilisaat member the kakilisaat member consists of silicic basalts, basaltic andesites and a few si-enriched picrites; mgo is in the range 6–13 wt%. the rocks have elevated sio2, k2o, rb, ba, th, u, light ree and pb, and low mgo, cao, tio2, feo* (total iron as feo), v, cu and ni relative to the rocks of the nunavik member (figs 72–75). these features are typical for rocks that are contaminated with continental crustal material, either crystalline basement or clastic sediments (e.g. larsen & pedersen 2009; pedersen et al. 2018). the occurrence of sediment xenoliths in the volcanic rocks suggests that the contamination took place in high-level magma chambers in the crust, presumably within the sediments of the nuussuaq basin. the low contents of v, cu, ni and co (fig. 74) indicate a possibility that the contamination process may have resulted in the formation at depth of mineral deposits with these elements. the relatively high nb contents (fig. 73) suggest that the original magmas had a somewhat enriched character akin to that of the nerutusoq member. the rocks of the kakilisaat member show compositional similarities to the crustally contaminated basalts of the tunoqqu and kûgánguaq members on nuussuaq and disko (larsen & pedersen 2009; pedersen et al. 2017). http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 85 of 121 mgo (wt% ) svartenhuk formation nunavik member enriched nunavik mb nerutusoq member kakilisaat member s io βo mgo (wt% ) al o mgo (wt% ) f eo* mgo (wt% ) cao mgo (wt% ) t io mgo (wt% ) na o . . . . . mgo (wt% ) k o βo . . . . . . mgo (wt% ) p o βo fig. 72 major-element variation diagrams for rocks of the vaigat formation. note the mgo-rich character of the nunavik member, the high cao and p₂o₅ in the nerutusoq member and the high sio₂ and low feo* and cao in the crustally contaminated kakilisaat member. βo1 is a marker horizon in the nunavik member. data points for the svartenhuk formation are shown in grey for comparison. data in wt% oxides. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 86 of 121 0 10 20 30 40 50 0 5 10 15 20 25 30 mgo (wt%) svartenhuk formation vaigat formation: nunavik mb enriched nunavik mb nerutusoq mb kakilisaat mb rb βo1 0 100 200 300 400 500 600 0 5 10 15 20 25 30 mgo (wt%) ba βo1 0 100 200 300 400 500 600 0 5 10 15 20 25 30 mgo (wt%) sr βo1 0 10 20 30 40 50 60 70 0 5 10 15 20 25 30 mgo (wt%) y 0 50 100 150 200 250 300 0 5 10 15 20 25 30 mgo (wt%) zr βo1 0 10 20 30 40 50 60 70 0 5 10 15 20 25 30 mgo (wt%) nb βo1 0 20 40 60 80 100 0 5 10 15 20 25 30 mgo (wt%) ce βo1 0 0.5 1 1.5 2 0 5 10 15 20 25 30 mgo (wt%) u βo1 fig. 73 incompatible trace-element variation diagrams for rocks of the vaigat formation. the nunavik member has low concentrations of these elements, but scattered lavas are enriched in sr, nb, ba and ce, akin to the nerutusoq member. the marker horizon βo1 has a unique, very enriched composition. the crustally contaminated kakilisaat member has unusually high nb, suggesting that some of these magmas are enriched in addition to being contaminated. data points for the svartenhuk formation are shown in grey for comparison. vertical axes in ppm. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 87 of 121 0 10 20 30 40 50 0 5 10 15 20 25 30 mgo (wt%) svartenhuk formation vaigat formation: nunavik mb enriched nunavik mb nerutusoq mb kakilisaat mb rb βo1 0 100 200 300 400 500 600 0 5 10 15 20 25 30 mgo (wt%) ba βo1 0 100 200 300 400 500 600 0 5 10 15 20 25 30 mgo (wt%) sr βo1 0 10 20 30 40 50 60 70 0 5 10 15 20 25 30 mgo (wt%) y 0 50 100 150 200 250 300 0 5 10 15 20 25 30 mgo (wt%) zr βo1 0 10 20 30 40 50 60 70 0 5 10 15 20 25 30 mgo (wt%) nb βo1 0 20 40 60 80 100 0 5 10 15 20 25 30 mgo (wt%) ce βo1 0 0.5 1 1.5 2 0 5 10 15 20 25 30 mgo (wt%) u βo1 mgo (wt% ) v mgo (wt% ) sc mgo (wt% ) cu mgo (wt% ) co mgo (wt% ) svartenhuk formation vaigat formation: nunavik mb enriched nunavik mb nerutusoq mb kakilisaat mb ni mgo (wt% ) cr fig. 74 transition-element variation diagrams for rocks of the vaigat formation. note the apparent loss of v, cu, co and ni in the kakilisaat member. data points for the svartenhuk formation are shown in grey for comparison. vertical axes in ppm. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 88 of 121 1 10 100 1000 la ce pr nd pm sm eu gd tb dy ho er tm yb lu vaigat formation nunavik member 1 10 100 1000 la ce pr nd pm sm eu gd tb dy ho er tm yb lu vaigat formation nerutusoq member 0 1 10 100 rb ba th u nb ta k la ce pb pr sr nd p sm zr hf eu ti gd tb dy y yb 0 1 10 100 rb ba th u nb ta k la ce pb pr sr nd p sm zr hf eu ti gd tb dy y yb 1 10 100 1000 la ce pr nd pm sm eu gd tb dy ho er tm yb lu vaigat formation kakilisaat member 0 1 10 100 rb ba th u nb ta k la ce pb pr sr nd p sm zr hf eu ti gd tb dy y yb fig. 75 ree and multi-element diagrams for representative rocks of the vaigat formation. different colours of curves in the nunavik member diagrams are intended to aid the distinction of the individual samples. the magenta samples in the nunavik member are incompatible-element-enriched; the two with the highest values are from the βo1 marker horizon. left diagrams: chondrite normalised. right diagrams: primitive mantle normalised. normalisation factors from mcdonough & sun (1995). http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 89 of 121 1 10 100 1000 la ce pr nd pm sm eu gd tb dy ho er tm yb lu vaigat formation nunavik member 1 10 100 1000 la ce pr nd pm sm eu gd tb dy ho er tm yb lu vaigat formation nerutusoq member 0 1 10 100 rb ba th u nb ta k la ce pb pr sr nd p sm zr hf eu ti gd tb dy y yb 0 1 10 100 rb ba th u nb ta k la ce pb pr sr nd p sm zr hf eu ti gd tb dy y yb 1 10 100 1000 la ce pr nd pm sm eu gd tb dy ho er tm yb lu vaigat formation kakilisaat member 0 1 10 100 rb ba th u nb ta k la ce pb pr sr nd p sm zr hf eu ti gd tb dy y yb table 5 chemical analyses of rocks from the vaigat formation member kakilisaat member nerutusoq member ggu no. 181015 181019 181022 181037 181073 181075 263818 181077 263822 263823 181013 263821 profile no. 4 4 4 13 13 48b 13 48b 48b 4 48b name kakilisaat kakilisaat kakilisaat saviit qaq. saviit s. saviit s. nerutusoq saviit s. nerutusoq nerutusoq kakilisaat nerutusoq major elements in wt% (xrf analyses) sio₂ 48.87 49.11 49.88 50.42 48.97 50.55 46.20 46.93 46.67 46.34 47.71 47.64 tio₂ 1.30 1.39 1.41 1.51 1.52 1.50 1.39 1.74 1.60 1.63 1.61 1.56 al₂o₃ 12.71 14.19 14.19 14.19 14.30 15.49 12.84 13.79 14.35 14.20 14.01 14.50 fe₂o₃ 4.42 3.22 2.39 2.62 3.38 1.42 3.23 4.73 4.58 4.89 4.63 3.34 feo 5.80 6.95 7.57 7.15 6.35 7.28 7.69 6.33 6.56 6.53 5.64 7.69 mno 0.20 0.24 0.20 0.20 0.20 0.20 0.22 0.29 0.17 0.17 0.20 0.19 mgo 11.46 9.39 8.63 8.19 6.09 5.69 11.29 8.98 8.20 7.82 7.64 7.62 cao 9.72 11.20 10.00 9.50 9.98 9.88 11.64 11.57 12.14 12.80 12.30 13.11 na₂o 1.39 2.11 1.79 2.61 3.32 2.16 1.68 1.83 1.79 1.78 1.78 1.88 k₂o 0.21 0.15 0.72 0.62 1.11 0.83 0.15 0.50 0.46 0.39 0.59 0.50 p₂o₅ 0.12 0.15 0.15 0.16 0.17 0.23 0.19 0.26 0.27 0.28 0.31 0.26 volatiles 3.03 1.78 3.05 2.53 3.53 3.95 3.65 2.54 3.25 2.93 2.12 1.70 sum 99.23 99.88 99.98 99.70 98.92 99.18 100.17 99.49 100.04 99.75 98.54 99.99 feo* 9.78 9.85 9.72 9.51 9.39 8.56 10.60 10.59 10.68 10.93 9.81 10.70 mg-no. 70.34 65.85 64.24 63.53 56.74 57.36 68.31 63.18 60.83 59.14 61.18 59.03 trace elements in ppm (icp-ms analyses) sc 33.1 34.6 34.1 36.4 35.3 32.3 32.7 38.3 37.5 35.9 36.7 36.3 v 243 260 260 280 304 219 311 324 333 337 283 335 cr 1134 436 405 418 88.1 269 572 471 271 260 411 243 co 51.8 47.8 47.0 44.4 40.2 32.8 58.2 56.6 51.7 48.9 42.8 52.8 ni 134 165 146 74.1 64.4 22.6 262 166 103 99.5 54.0 77.86 cu 67.6 98.2 93.8 39.7 86.0 21.5 122 85.3 139 139 39.7 107 zn 73.5 95.9 73.6 80.7 68.5 76.1 89.0 71.4 88.9 88.6 75.4 76.0 ga 17.7 18.2 18.4 18.1 17.8 21.3 21.0 19.1 21.5 22.0 20.7 24.0 rb 2.64 1.16 12.3 13.6 33.2 33.5 2.44 13.0 13.4 10.3 6.81 12.1 sr 205 264 298 269 346 251 179 299 276 305 452 314 y 20.2 22.1 21.9 25.7 24.0 26.3 21.5 24.3 23.7 24.4 25.5 23.0 zr 89.0 105 101 127 113 150 70.1 125 117 119 143 115 nb 6.45 11.7 10.7 10.0 15.8 27.13 4.80 38.3 17.7 17.4 31.3 17.8 cs 0.009 0.027 0.157 0.067 0.119 0.773 0.041 0.102 0.041 0.193 0.033 0.218 ba 120 154 325 298 984 354 51.5 328 144 146 533 211 la 10.0 14.2 13.1 15.3 22.4 30.2 5.15 34.6 17.1 17.4 50.4 17.3 ce 21.6 30.9 28.7 32.5 50.2 61.6 13.0 67.8 37.0 37.8 92.1 38.8 pr 2.90 4.09 3.95 4.28 6.62 7.51 2.04 7.82 4.88 5.07 10.6 5.09 nd 12.5 17.5 17.0 18.0 26.9 29.5 10.2 28.6 20.3 21.4 38.3 21.2 sm 3.12 4.10 4.10 4.41 5.40 5.73 3.02 5.26 4.54 4.59 6.45 4.66 eu 1.06 1.36 1.33 1.39 1.68 1.72 1.09 1.61 1.49 1.49 1.90 1.49 gd 3.65 4.31 4.33 4.62 5.35 5.69 3.59 5.30 4.74 4.77 6.06 4.79 tb 0.598 0.692 0.675 0.788 0.815 0.891 0.622 0.807 0.747 0.777 0.932 0.766 dy 3.50 3.84 3.90 4.31 4.37 4.76 3.70 4.51 4.31 4.49 4.90 4.28 ho 0.717 0.769 0.787 0.880 0.897 0.947 0.771 0.884 0.886 0.879 0.902 0.862 er 1.97 2.12 2.10 2.43 2.39 2.51 2.07 2.34 2.28 2.40 2.46 2.30 tm 0.284 0.292 0.305 0.342 0.357 0.349 0.283 0.347 0.320 0.346 0.348 0.323 yb 1.79 1.87 1.87 2.20 2.08 2.24 1.80 2.06 2.00 2.04 2.06 2.08 lu 0.268 0.286 0.286 0.312 0.312 0.342 0.256 0.304 0.276 0.287 0.302 0.296 hf 2.39 2.59 2.78 3.22 3.05 3.94 1.98 3.38 2.93 3.01 3.49 3.10 ta 0.384 0.522 0.531 0.478 0.424 1.37 0.277 2.17 1.27 1.26 1.56 1.15 pb 1.87 3.85 3.49 3.62 7.42 3.81 1.19 2.19 1.88 2.11 3.95 2.21 th 1.66 2.28 2.15 2.31 2.74 4.60 0.396 3.09 1.44 1.49 3.02 1.60 u 0.379 0.299 0.296 0.378 0.238 1.007 0.130 0.934 0.477 0.455 0.959 0.481 feo*= total iron calculated as feo. mg-number = 100 × atomic mg/(mg+fe²⁺), with the iron oxidation ratio adjusted to fe₂o₃/feo = 0.15. for notes on the samples, see next pages. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 90 of 121 table 5 (continued) chemical analyses of rocks from the vaigat formation member nunavik member ggu no. 181066 181067 251515 251525 251529 278353 181085 251589 278354 251527 251588 181056 profile no. 24 24 9 9 9 44c 14 30 44c 9 30 24 name pk 1120 pk 1120 pk 1078 pk 1078 pk 1078 kuugartorf. saviit n. iviss. itinn. kuugartorf. pk 1078 iviss.itinn. pk 1120 major elements in wt% (xrf analyses) sio₂ 42.73 43.18 45.55 45.10 45.21 45.92 45.16 45.87 45.68 43.16 46.46 47.40 tio₂ 0.79 0.91 1.45 1.23 1.30 1.24 1.39 1.68 1.35 1.68 1.59 1.92 al₂o₃ 7.34 8.48 9.44 10.41 10.50 10.98 11.33 11.26 12.54 11.91 13.20 13.24 fe₂o₃ 2.38 3.00 3.59 3.52 2.15 2.49 4.93 4.43 5.05 6.28 3.48 6.60 feo 9.08 7.99 8.18 8.32 9.49 8.63 7.03 7.48 6.77 6.82 8.10 5.04 mno 0.22 0.21 0.17 0.17 0.17 0.17 0.22 0.17 0.19 0.22 0.17 0.23 mgo 26.42 24.14 20.41 19.19 18.39 17.95 15.77 14.94 12.40 11.57 10.10 9.51 cao 6.21 6.93 7.98 8.90 8.76 8.58 9.53 9.45 10.01 11.63 11.29 11.94 na₂o 0.60 0.89 1.48 1.44 1.73 1.29 1.37 1.62 1.36 0.92 1.70 1.82 k₂o 0.28 0.06 0.06 0.06 0.10 0.13 0.11 0.15 0.07 0.64 0.28 0.24 p₂o₅ 0.07 0.09 0.13 0.14 0.14 0.12 0.11 0.15 0.12 0.25 0.16 0.18 volatiles 3.78 3.68 1.76 1.71 2.00 2.32 2.70 2.76 4.59 5.51 3.73 1.72 sum 99.89 99.55 100.20 100.19 99.94 99.82 99.65 99.96 100.13 100.59 100.26 99.84 feo* 11.22 10.69 11.41 11.49 11.42 10.87 11.47 11.47 11.31 12.47 11.23 10.98 mg-no. 82.64 82.04 78.35 77.17 76.50 76.96 73.56 72.49 68.91 65.25 64.53 63.66 trace elements in ppm (icp-ms analyses) sc 23.4 26.1 26.3 30.7 28.7 29.3 33.2 31.2 34.9 35.5 34.5 38.3 v 178 211 271 283 282 262 284 306 306 324 326 329 cr 1834 1499 1287 1370 1287 1258 1214 971 686 711 505 358 co 101 91.7 79.4 80.6 76.5 70.0 78.5 64.7 62.6 69.5 55.6 53.3 ni 1343 1225 937 844 817 737 543 529 341 289 236 195 cu 70.2 43.9 86.3 121 100 80.3 129 135 30.1 118 131 169 zn 74.3 76.1 81.5 88.3 86.0 77.4 82.1 86.3 83.7 91.3 82.1 88.1 ga 8.1 11.9 15.2 16.6 16.8 15.5 16.9 17.0 16.8 17.6 19.3 19.8 rb 6.21 2.17 0.29 0.77 1.45 1.52 1.49 1.15 1.27 24.6 3.75 2.24 sr 107 117 161 129 183 158 163 194 151 356 279 237 y 10.9 13.8 18.9 18.3 18.5 19.4 20.4 22.0 21.1 20.3 21.9 25.4 zr 39.5 44.6 84.8 65.2 68.2 74.7 76.1 98.6 74.3 128 94.0 106 nb 1.78 1.67 5.29 2.80 3.53 5.85 4.10 5.59 4.56 36.2 13.3 10.6 cs 0.013 0.016 0.005 0.037 0.062 0.042 0.013 0.025 0.046 0.114 0.011 0.005 ba 25.6 23.9 24.9 19.5 27.4 57.9 24.7 42.2 32.1 230 82.3 62.7 la 1.96 1.80 5.05 2.98 3.34 5.98 4.15 5.66 4.87 19.2 9.71 8.14 ce 5.61 5.47 13.6 8.74 9.60 14.6 11.2 15.4 12.4 41.6 23.0 20.3 pr 0.96 0.96 2.24 1.53 1.62 2.22 1.78 2.45 1.95 5.20 3.21 3.11 nd 5.12 5.38 11.1 8.04 8.30 10.3 9.21 12.5 9.67 20.5 14.4 14.7 sm 1.55 1.78 3.34 2.65 2.75 2.97 2.92 3.80 2.93 4.18 3.77 4.10 eu 0.578 0.659 1.10 0.905 1.02 1.06 1.03 1.34 1.06 1.32 1.31 1.46 gd 1.78 2.22 3.69 3.03 3.25 3.49 3.25 4.21 3.59 4.38 4.24 4.60 tb 0.310 0.380 0.580 0.508 0.559 0.577 0.570 0.697 0.586 0.668 0.689 0.770 dy 1.77 2.30 3.49 3.14 3.29 3.39 3.42 3.93 3.54 3.67 3.99 4.48 ho 0.383 0.493 0.696 0.642 0.636 0.679 0.698 0.802 0.723 0.720 0.796 0.902 er 1.03 1.33 1.83 1.67 1.79 1.84 1.84 2.08 1.95 1.86 2.09 2.40 tm 0.151 0.202 0.259 0.258 0.262 0.266 0.289 0.290 0.285 0.277 0.277 0.338 yb 0.90 1.17 1.54 1.52 1.58 1.61 1.68 1.64 1.72 1.67 1.72 2.09 lu 0.138 0.185 0.226 0.226 0.230 0.258 0.246 0.260 0.264 0.243 0.265 0.307 hf 1.14 1.30 2.19 1.72 1.86 1.94 1.99 2.49 1.93 3.28 2.54 2.86 ta 0.190 0.116 0.395 0.204 0.292 0.371 0.339 0.411 0.284 2.33 0.866 0.689 pb 0.180 0.185 0.581 0.348 0.345 1.09 0.475 0.604 0.722 0.672 0.633 0.662 th 0.121 0.109 0.406 0.187 0.178 0.689 0.276 0.358 0.366 2.520 0.806 0.758 u 0.035 0.033 0.111 0.055 0.064 0.203 0.088 0.103 0.108 0.551 0.216 0.224 87sr/86sr 0.703473 0.703340 0.703310 0.703232 0.703275 143nd/144nd 0.513081 0.513103 0.513045 0.513082 0.513070 sr-nd isotope data (as measured) from holm et al. (1993). http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 91 of 121 table 5 (continued) notes on analysed samples from the vaigat formation 181015 fine-grained olivine-phyric magnesian basalt, slightly crustally contaminated. thick, compound pahoehoe flow, uppermost flow in the kakilisaat member. profile 4, kakilisaat, south-east svartenhuk halvø. 181019 fine-grained olivine-phyric magnesian basalt, crustally contaminated. compound pahoehoe flow beneath the uppermost hyaloclastite horizon in the kakilisaat member. profile 4, kakilisaat, south-east svartenhuk halvø. 181022 fine-grained aphyric silicic basalt, crustally contaminated. pillow lava horizon on top of 75 m foreset-bedded hyaloclastite low in the kakilisaat member. profile 4, kakilisaat, south-east svartenhuk halvø. 181037 fine-grained aphyric silicic basalt, crustally contaminated. lava flow in the upper kakilisaat member, western corner of the mouth of ulissat qooruat, south-east coast of svartenhuk halvø. located on fig. 17. 181073 fine-grained aphyric silicic basalt, crustally contaminated. kakilisaat member lava flow low in profile 13, saviit south, southern svartenhuk halvø. 181075 plagioclase-olivine-phyric basaltic andesite, crustally contaminated and enriched in incompatible trace elements (th, u, nb, ree). kakilisaat member lava flow in profile 13, saviit south, southern svartenhuk halvø. 263818 aphyric magnesian basalt. brown hyaloclastite, nerutusoq member in profile 48b, nerutusoq valley at 360 m altitude, northern svartenhuk halvø. 181077 aphyric basalt distinctly enriched in incompatible trace elements (ba, th, u, nb, ree). lava flow, nerutusoq member in profile 13, saviit south, southern svartenhuk halvø. 263822 aphyric basalt enriched in incompatible trace elements (ba, th, u, nb, ree). brown pillow breccia, nerutusoq member in profile 48b, nerutusoq valley, northern svartenhuk halvø. 263823 aphyric basalt enriched in incompatible trace elements (ba, th, u, nb, ree). brown lava flow, nerutusoq member in profile 48b, nerutusoq valley, northern svartenhuk halvø. 181013 fine-grained aphyric basalt distinctly enriched in incompatible trace elements (ba, th, u, nb, ree). pahoehoe lava flow, nerutusoq member in profile 4, kakilisaat, south-east svartenhuk halvø. 263821 aphyric basalt enriched in incompatible trace elements (ba, th, u, nb, ree). black hyaloclastite, nerutusoq member in profile 48b, nerutusoq valley, northern svartenhuk halvø. 181066 olivine-rich picrite. pahoehoe flow lobe in grey compound lava flow, lower nunavik member in profile 24, peak 1120 m, southern svartenhuk halvø. 181067 olivine-rich picrite. pahoehoe flow lobe in grey compound lava flow, lower nunavik member in profile 24, peak 1120 m, southern svartenhuk halvø. 251515 olivine-rich picrite. pahoehoe flow lobe in grey compound lava flow, upper nunavik member in profile 9, peak 1078 m, eastern svartenhuk halvø. 251525 olivine-microphyric picrite. pahoehoe flow lobe in grey compound lava flow, middle nunavik member in profile 9, peak 1078 m, eastern svartenhuk halvø. 251529 medium-grained picrite. pahoehoe flow lobe in grey compound lava flow, lower nunavik member in profile 9, peak 1078 m, eastern svartenhuk halvø. 278353 picrite. pahoehoe flow lobe in greenish grey, compound lava flow, upper nunavik member in profile 44c, kuugaartorfik, northern svartenhuk halvø. 181085 olivine-poor picrite. sheet flow 25 m in thickness, middle nunavik member in profile 14, saviit north, southern svartenhuk halvø. 251589 medium-grained picrite. pahoehoe flow lobe in grey compound lava flow, upper nunavik member in profile 30, ivissukkat itinnerat, eastcentral svartenhuk halvø. 278354 mg-poor picrite. pahoehoe flow lobe in crumbling compound lava flow, upper nunavik member in profile 44c, kuugaartorfik, northern svartenhuk halvø. 251527 olivine-microphyric magnesian basalt distinctly enriched in incompatible trace elements (ba, th, u, nb, zr, sr, ree). pahoehoe flow lobe at base of dark marker horizon βo1. compound lava flow, middle nunavik member in profile 9, peak 1078 m, eastern svartenhuk halvø. 251588 medium-grained magnesian basalt. pahoehoe flow lobe in grey compound lava flow, upper nunavik member in profile 30, ivissukkat itinnerat, east-central svartenhuk halvø. 181056 olivine-phyric basalt. lava flow with pseudopillow structure, upper nunavik member in profile 24, peak 1120 m, southern svartenhuk halvø. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 92 of 121 svartenhuk formation the volcanic rocks of the svartenhuk formation comprise tholeiitic basalts, subordinate picrites and a few basaltic andesites. they have a total range of 3.4–21.7 wt% mgo with an average of 7.2 wt% mgo (620 analyses). a few basalts are crustally contaminated and some scattered basalts are geochemically enriched. major element variation diagrams for the svartenhuk formation are shown in fig. 76, trace-element variation diagrams in figs 77 and 78, and ree and multi-element diagrams in fig. 79. note that figs 76–78 use the mg-number as abscissa and not the mgo content as in the plots of the vaigat formation; this is because the major part of the basalts have a limited variation in mgo content (5–9 wt% mgo), which does not produce a sufficiently large spread on the abscissa. the figures are designed to be comparable with the similar geochemistry diagrams for the maligât formation in pedersen et al. (2018). representative chemical analyses are shown in table 6. tunuarsuk member the tunuarsuk member comprises basalts with 5–10 wt% mgo grading into magnesian basalts with 10–12 wt% mgo and scattered picrites with 12–16 wt% mgo; a few rocks with up to 22 wt% mgo probably contain accumulated olivine. the average mgo content is 8.4 wt%, and the majority of the rocks are olivine normative. tio2 is in the range 1.1–3.3 wt% with an average of 1.8 wt% tio2, and with few exceptions the p2o5/tio2 ratio is less than 0.11. the magmas were relatively depleted in incompatible trace elements such as rb, ba, th, u, nb, and light ree (fig. 79). some flows scattered throughout the region are distinctly enriched in these same elements, and very few flows (2 analyses) are crustally contaminated. many ree and multi-element curves in fig. 79 cross each other. nuuit member the nuuit member comprises basalts with 3.4–11 wt% mgo; only 3 out of 220 analyses have more than 9 wt% mgo and the average mgo content is 6.4 wt%. tio2 is in the range 1.5–3.8 wt%, and the average tio2 content is 2.4 wt%. the basalts are similar in composition to the more ti-rich (tio2 >2.0 wt%) basalts of the tunuarsuk member. the magmas were relatively depleted in incompatible trace elements such as rb, ba, th, u, nb and light ree, similar to the basalts of the tunuarsuk member. some flows are enriched in the same elements, as discussed below. with the exception of these, the ree and multi-element curves in fig. 79 are near-parallel and do not cross each other. this is in contrast to the same curves for the tunuarsuk member. a few scattered flows are crustally contaminated (4 analyses have sio2 ≥51 wt%) and some more may be slightly contaminated. skalø member the skalø member comprises basalts with 4.5–10 wt% mgo; only 3 out of 127 analyses have more than 9 wt% mgo and the average mgo content is 6.5 wt%. tio2 is in the range 1.1–3.5 wt%, and the average tio2 content is 2.0 wt%. these compositions are very similar to those of the nuuit member, with slightly lower tio2 in the skalø member. sio2, al2o3 and cao range to slightly higher values than in the nuuit member, and many skalø member flows probably have relatively high contents of plagioclase. the magmas were relatively depleted in incompatible trace elements such as rb, ba, th, u, nb and light ree, similar to the basalts of the tunuarsuk and nuuit members. a single flow is enriched in the same elements (fig. 79). a few flows may be crustally contaminated; 6 analyses have sio2 slightly higher than 51 wt% but still form part of the main compositional trends. with some exceptions the ree and multi-element curves in fig. 79 are near-parallel and in this respect more similar to the curves of the nuuit member than to the tunuarsuk member. the skalø member flows at aputituut (profile 55) have relatively evolved compositions with high tio2 and p2o5. the uppermost flow is extreme, with 4.0 wt% tio2 and 0.47 wt% p2o5 at 5.7 wt% mgo. these values are similar to those in the overlying arfertuarsuk member. however, the trace-element ratios (ree and zr-nb-y) clearly indicate that all the flows belong to the skalø member and the top flow is just highly evolved. discussion the three volcanic members of the svartenhuk formation are compositionally very similar in their major and trace elements, and what defines the chemostratigraphy is the pattern of variation up-section. for example, the major difference between the skalø and nuuit members is the variable tio2 compositional profiles in the nuuit member compared to the near-constant profiles in the skalø member (fig. 42). the basalts of the svartenhuk formation represent fig. 76 (figure on next page) major-element variation diagrams for rocks of the svartenhuk formation. crustally contaminated and enriched rocks are included in the main members because their major elements are as a rule not distinguishable from those of the main members. a few exceptions are labelled e: enriched, and c: crustally contaminated. note that the abscissa is the mg-number (atomic 100×mgo/(mgo+feo)), not the simple mgo as in the vaigat formation plots. the skalø member partly conceals the nuuit and tunuarsuk members because it is plotted on top of them. fractionation goes from right to left. the fractionation trends bend at mg-numbers around 57 (7–8 wt% mgo) because at that stage the fractionating olivine is joined by plagioclase and shortly afterwards by clinopyroxene. vertical axes in wt% oxides. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 93 of 121 40 45 50 55 60 30 35 40 45 50 55 60 65 70 75 80 85 mg -number nuuit member skalø membersio 2 e cc 5 10 15 20 30 35 40 45 50 55 60 65 70 75 80 85 mg -number al2o3 5 10 15 20 30 35 40 45 50 55 60 65 70 75 80 85 mg -number feo* e c 0 5 10 15 30 35 40 45 50 55 60 65 70 75 80 85 mg -number cao e e c c 0 1 2 3 4 5 30 35 40 45 50 55 60 65 70 75 80 85 mg -number tio 2 e e c 0 1 2 3 4 5 30 35 40 45 50 55 60 65 70 75 80 85 mg -number na2o e c e 0.0 0.5 1.0 1.5 2.0 30 35 40 45 50 55 60 65 70 75 80 85 mg -number k 2o e e c 0.0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 30 35 40 45 50 55 60 65 70 75 80 85 mg -number p 2o5 e e c tunuarsuk member fig. 76 (caption on previous page) http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 94 of 121 mg-number r b mg-number ba mg-number sr mg-number y mg-number zr mg-number nb mg-number ce . . mg-number u dykes and sills arfertuarsuk mb skalø mb nuuit mb tunuarsuk mb circles: enriched svartenhuk fm fig. 77 incompatible trace-element variation diagrams for rocks of the svartenhuk and naqerloq formations and intrusions. the tunuarsuk, nuuit and skalø members have similar low element concentrations, but scattered lavas in all three members are enriched (open circles). the arfertuarsuk member has significantly higher concentrations of these elements. intrusions compositionally similar to both formations are present. vertical axes in ppm. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 95 of 121 mg -number v mg -number sc mg -number cu mg -number zn mg -number ni four samples off scale at – ppm mg -number cr five samples off scale at ppm – dykes and sills arfertuarsuk mb skalø mb nuuit mb tunuarsuk mb circles: enriched svartenhuk fm fig. 78 transition-element variation diagrams for rocks of the svartenhuk and naqerloq formations and intrusions. vertical axes in ppm. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 96 of 121 la c e p r nd p m s m e u gd t b dy ho e r t m y b lu r b ba t h u nb t a k la c e p b p r s r nd p s m zr hf e u t i gd t b dy y y b la c e p r nd p m s m e u gd t b dy ho e r t m y b lu r b ba t h u nb t a k la c e p b p r s r nd p s m zr hf e u t i gd t b dy y y b la c e p r nd p m s m e u gd t b dy ho e r t m y b lu r b ba t h u nb t a k la c e p b p r s r nd p s m zr hf e u t i gd t b dy y y b svartenhuk formation skalø member svartenhuk formation nuuit member svartenhuk formation tunarsuk member svartenhuk formation fig. 79 ree and multi-element diagrams for representative rocks of the svartenhuk formation. the different colours of the curves are intended to aid the distinction of the individual samples. the magenta samples are incompatible-element-enriched. note the particularly variable slope of the la–sm limb of the ree curve for the tunuarsuk member. left diagrams: chondrite normalised. right diagrams: primitive mantle normalised. normalisation factors from mcdonough & sun (1995). http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 97 of 121 table 6 chemical analyses of rocks from the svartenhuk formation member tunuarsuk member nuuit member ggu no. 181120 278559 251317 251281 262803 251295 278343 251333 262815 262793 566511 566505 profile no. 50 42 38 64b 38 44c 37 64b 64a 65 reconn. name qinnivik umiiarfik marraarn. qaqqap q. paannivik qaqqap q. kuugaart.nuup qaa. paannivik paannivik salliaruseq e nunatak major elements in wt% (xrf analyses) sio₂ 45.19 45.98 45.89 47.74 47.62 48.39 47.92 48.16 47.89 49.08 48.07 48.66 tio₂ 1.18 1.26 1.16 1.55 1.64 1.86 2.06 1.92 2.78 2.26 1.69 1.47 al₂o₃ 11.18 11.86 12.78 13.09 12.70 14.24 14.02 14.19 13.63 13.88 14.56 14.87 fe₂o₃ 2.56 5.60 5.18 12.15* 4.88 12.70* 4.07 5.46 5.05 4.76 12.28* 11.79* feo 7.75 6.14 6.05 7.16 8.15 6.72 8.35 8.19 mno 0.21 0.18 0.17 0.19 0.18 0.191 0.18 0.18 0.21 0.23 0.19 0.19 mgo 16.70 15.69 14.43 11.83 10.29 7.34 7.20 6.72 6.20 5.62 8.14 8.09 cao 9.44 9.47 9.54 10.52 11.13 11.88 11.77 12.57 11.04 10.70 11.80 12.37 na₂o 1.39 1.63 1.61 1.93 1.87 2.22 2.40 2.47 2.61 2.63 1.98 1.93 k₂o 0.50 0.09 0.12 0.18 0.19 0.21 0.24 0.28 0.31 0.19 0.111 0.164 p₂o₅ 0.10 0.12 0.10 0.15 0.16 0.17 0.21 0.25 0.31 0.23 0.135 0.123 volatiles 3.25 2.10 2.92 0.41 2.27 0.36 1.68 1.36 1.51 2.41 0.80 -0.02 sum 99.45 100.12 99.95 99.73 100.08 99.56 99.90 100.28 99.89 100.17 99.75 99.64 feo* 10.05 11.18 10.71 10.93 11.55 11.43 11.81 11.63 12.89 12.47 11.05 10.61 mg-no. 77.06 73.95 73.15 68.65 64.30 56.51 55.22 53.89 49.30 47.68 59.85 60.67 trace elements in ppm (icp-ms analyses) sc 32.6 28.7 32.8 35.2 35.7 37.1 35.1 37.0 34.0 38.1 42.8 44.2 v 248 254 264 339 314 360 355 361 418 427 356 337 cr 1134 951 900 529 617 315 273 161 122 86.4 343 353 co 70.1 65.7 65.2 55.5 59.2 48.7 48.0 46.8 48.3 45.1 49.5 49.4 ni 645 531 462 401 280 123 113 93.8 73.9 64.9 154 136 cu 157 113 125 164 141 172 170 177 196 251 211 178 zn 70.7 74.9 75.3 80.6 89.0 90.6 96.7 94.4 119 116 90.1 84.9 ga 14.9 15.8 16.1 18.5 17.9 20.5 21.7 22.0 23.0 22.1 19.7 18.9 rb 12.4 1.01 1.25 1.90 3.36 1.60 2.22 4.13 4.76 4.45 0.63 2.00 sr 140 181 136 177 194 216 281 361 307 188 167 175 y 19.4 20.3 20.3 27.5 24.9 28.9 26.8 29.2 34.1 40.1 28.6 25.6 zr 62.7 71.4 65.1 88.9 88.4 102 114 117 162 143 91.3 81.0 nb 3.41 10.1 4.47 7.87 6.13 6.39 7.34 24.4 12.2 5.28 4.02 4.30 cs 0.043 0.080 0.014 0.026 0.150 0.012 0.032 0.069 0.054 0.071 0.007 0.006 ba 27.3 70.0 37.0 67.2 66.8 47.5 53.8 87.1 93.1 54.1 25.0 30.7 la 3.47 11.4 4.89 8.46 8.43 6.91 7.29 17.8 12.0 7.14 4.54 4.34 ce 9.18 21.7 11.6 18.6 19.2 17.0 19.2 39.4 30.1 18.8 12.8 11.9 pr 1.50 2.78 1.73 2.62 2.68 2.58 3.05 5.25 4.62 3.05 2.14 1.92 nd 7.98 12.0 8.63 12.1 12.3 12.9 15.0 21.6 21.8 15.5 11.3 10.2 sm 2.49 3.09 2.70 3.50 3.54 4.02 4.55 5.03 6.12 5.21 3.70 3.31 eu 0.901 1.03 0.955 1.25 1.22 1.41 1.60 1.67 2.07 1.79 1.37 1.23 gd 3.00 3.56 3.13 4.29 4.28 4.62 5.06 5.53 7.03 6.40 4.68 4.09 tb 0.511 0.606 0.552 0.751 0.667 0.808 0.841 0.890 1.10 1.09 0.775 0.712 dy 3.32 3.51 3.49 4.71 4.15 4.71 4.88 5.19 6.10 6.58 5.10 4.42 ho 0.677 0.724 0.723 0.963 0.837 1.00 0.941 1.04 1.21 1.41 1.03 0.922 er 1.84 1.91 1.94 2.61 2.15 2.53 2.52 2.69 3.12 3.70 2.77 2.51 tm 0.273 0.268 0.287 0.375 0.328 0.372 0.340 0.387 0.459 0.585 0.426 0.378 yb 1.66 1.63 1.74 2.41 1.90 2.27 2.07 2.38 2.69 3.34 2.52 2.26 lu 0.255 0.259 0.259 0.362 0.289 0.342 0.314 0.349 0.402 0.500 0.376 0.344 hf 1.77 1.87 1.78 2.36 2.38 2.75 2.88 2.97 4.19 3.71 2.53 2.22 ta 0.403 0.454 0.295 0.511 0.345 0.567 0.487 1.65 0.800 0.363 0.318 0.585 pb 0.374 0.497 0.452 0.644 0.727 0.643 0.769 0.814 1.30 1.06 0.490 0.499 th 0.249 0.914 0.296 0.560 0.467 0.448 0.580 2.00 1.01 0.611 0.306 0.334 u 0.074 0.233 0.102 0.159 0.141 0.142 0.164 1.06 0.283 0.191 0.087 0.112 87sr/86sr 0.703407 0.703650 0.703611 143nd/144nd 0.513045 0.512832 0.512974 feo*= total iron calculated as feo. total iron determined as fe₂o₃ is marked by an * mg-number = 100 × atomic mg/(mg+fe²+), with the iron oxidation ratio adjusted to fe₂o₃/feo = 0.15. for notes on the samples, see next page. sr-nd isotope data (as measured) from holm et al. (1993). http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 98 of 121 table 6 (continued) chemical analyses of rocks from the svartenhuk formation unit nuuit member skalø member ggu no. 263525 181097 181102 566520 262973 251338 278520 262896 262754 262988 262912 262697 profile no. 62 40 40 65 62 37 50 41b 64a 62 41b 55 name sigguk pann.qaq. pann.qaq. salliaruseq sigguk nuup q. umiiarfik skalø paannivik sigguk skalø aputituut major elements in wt% (xrf analyses) sio₂ 48.40 47.57 48.74 48.45 47.28 48.52 47.04 48.16 48.54 48.80 48.25 47.27 tio₂ 1.61 2.99 2.79 2.29 3.39 2.68 1.38 1.50 1.87 2.19 2.78 3.90 al₂o₃ 14.96 12.79 12.99 13.49 13.30 13.94 14.18 15.12 15.23 14.87 13.20 12.77 fe₂o₃ 6.58 6.90 3.08 14.81* 6.53 8.33 4.98 4.43 4.84 4.55 4.25 5.58 feo 5.00 7.87 10.03 8.78 6.54 6.11 6.23 6.32 8.56 9.90 8.92 mno 0.20 0.28 0.26 0.23 0.26 0.20 0.17 0.18 0.16 0.23 0.24 0.24 mgo 7.55 6.42 6.30 6.30 5.99 5.11 9.79 8.08 6.37 5.94 5.73 5.54 cao 12.19 10.89 10.89 11.06 10.45 10.57 11.64 12.06 11.98 11.38 10.84 10.52 na₂o 2.25 2.40 2.60 2.22 1.78 2.80 1.78 2.01 2.39 2.53 2.53 2.79 k₂o 0.21 0.22 0.34 0.24 0.18 0.34 0.42 0.14 0.11 0.14 0.33 0.52 p₂o₅ 0.15 0.30 0.31 0.22 0.36 0.32 0.11 0.13 0.17 0.21 0.29 0.46 volatiles 1.35 1.19 1.14 0.33 2.08 1.15 2.59 1.99 2.02 1.16 1.66 2.01 sum 100.45 99.82 99.47 99.64 100.38 100.50 100.19 100.03 100.00 100.56 100.00 100.52 feo* 10.92 14.08 12.80 13.32 14.66 14.04 10.59 10.22 10.67 12.65 13.72 13.94 mg-no. 58.31 47.99 49.89 48.89 45.26 42.42 65.16 61.54 54.70 48.72 45.80 44.58 trace elements in ppm (icp-ms analyses) sc 41.2 42.9 39.2 40.4 45.0 36.0 38.4 39.4 38.0 39.5 42.8 41.2 v 330 473 452 411 552 424 304 323 369 412 513 621 cr 341 157 171 135 115 31.6 447 427 277 75.3 92.9 74.7 co 44.9 49.7 46.5 47.7 45.9 46.1 51.0 47.1 46.6 48.1 50.7 59.0 ni 121 73.0 82.4 78.5 66.8 57.0 199 146 156 62.1 58.8 60.8 cu 196 326 284 224 442 292 167 174 200 269 360 367 zn 81.0 129 125 117 143 122 73.5 77.9 85.7 106 126 140 ga 20.0 22.7 24.0 21.4 23.3 24.7 18.3 18.9 20.5 21.5 22.3 24.7 rb 1.51 1.37 8.31 4.99 2.82 7.41 6.42 1.00 0.49 1.11 6.09 12.08 sr 181 195 246 194 191 304 206 174 200 197 200 236 y 27.8 48.1 44.1 37.8 58.2 46.2 22.8 26.2 30.3 37.5 46.3 58.4 zr 90.9 180 175 138 243 178 71.8 82.7 100 133 180 258 nb 5.30 12.2 21.9 10.7 15.2 26.3 5.00 4.21 5.45 7.83 12.4 20.7 cs 0.005 0.020 0.085 0.063 0.073 0.141 0.062 0.016 0.001 0.011 0.069 0.143 ba 45.6 67.7 148 61.3 60.5 190 42.9 30.9 29.6 45.8 86.5 140 la 5.00 10.3 21.3 9.82 15.2 26.9 5.31 4.50 5.89 7.93 12.8 18.2 ce 13.4 26.6 43.1 24.0 37.4 52.8 13.1 12.1 16.0 20.2 30.8 45.4 pr 2.19 4.25 5.97 3.55 5.74 6.79 2.04 2.00 2.52 3.19 4.59 6.97 nd 11.2 21.7 26.6 17.6 28.8 28.4 10.1 10.3 13.1 16.2 22.0 33.0 sm 3.54 6.65 6.86 5.27 8.15 6.84 3.05 3.30 4.03 4.92 6.46 9.20 eu 1.32 2.20 2.18 1.80 2.57 2.21 1.11 1.24 1.45 1.73 2.12 2.90 gd 4.35 7.71 7.69 6.47 9.68 7.80 3.60 4.10 5.12 5.82 7.65 10.59 tb 0.734 1.34 1.29 1.08 1.63 1.31 0.629 0.722 0.867 1.03 1.32 1.73 dy 4.71 8.24 7.89 6.68 9.97 7.76 3.76 4.45 5.43 6.39 7.67 10.26 ho 0.957 1.72 1.63 1.35 2.12 1.61 0.817 0.912 1.13 1.34 1.66 2.08 er 2.62 4.67 4.38 3.64 5.71 4.28 2.21 2.44 3.01 3.53 4.44 5.53 tm 0.399 0.681 0.630 0.533 0.857 0.629 0.324 0.382 0.444 0.556 0.654 0.834 yb 2.42 4.33 3.89 3.31 5.23 3.92 1.91 2.29 2.69 3.31 3.95 4.96 lu 0.339 0.609 0.566 0.506 0.801 0.565 0.287 0.346 0.415 0.511 0.590 0.770 hf 2.44 4.74 4.84 3.64 6.33 4.48 1.86 2.26 2.84 3.65 4.53 6.41 ta 0.316 0.780 1.33 0.844 0.846 1.29 0.286 0.256 0.342 0.444 0.784 1.29 pb 0.502 1.13 1.55 0.825 1.71 1.59 0.501 0.501 0.676 0.841 1.23 1.74 th 0.398 0.900 1.99 0.988 1.66 2.53 0.424 0.303 0.449 0.723 1.08 1.67 u 0.119 0.215 0.573 0.290 0.491 0.630 0.122 0.100 0.091 0.188 0.330 0.509 http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 99 of 121 table 6 (continued) notes on analysed samples from the svartenhuk formation 181120 picrite with platy and equant olivine phenocrysts. single olivin-rich lava flow within a succession of plagioclase glomerophyric and aphyric basalts. tunuarsuk member, qinnivik peninsula 1 km wsw of point 631 m, south-western svartenhuk halvø, indicated on the geological map (larsen 1983) as βo. 278559 aphyric picrite distinctly enriched in incompatible trace elements (ba, th, u, nb, sr, ree). pahoehoe flow lobe in compound flow c. 10 m thick and with 1 m dark red scoriaceous top zone. tunuarsuk member in profile 50, umiiarfik south, northern svartenhuk halvø. 251317 medium-grained picrite. pahoehoe flow lobe, part of dark greenish compound flow, uppermost tunuarsuk member in profile 42, marrarnaq, northern qeqertaq peninsula 11 km south-east of the settlement kangersuatsiaq (prøven). 251281 olivine-phyric magnesian basalt. massive lava flow 50 m thick and with 10 m scoriaceous top zone, uppermost flow in tunuarsuk member in profile 38, qaqqap qaa, east coast of the innerit peninsula. 262803 plagioclase-phyric magnesian basalt. massive lava flow c. 40 m thick, middle tunuarsuk member in profile 64, paannivik, southern innerit. 251295 plagioclase-phyric basalt. massive lava flow c. 40 m thick, middle tunuarsuk member in profile 38, qaqqap qaa, east coast of the innerit peninsula. 278343 aphyric basalt. massive lava flow c. 20 m thick, between hyaloclastite horizons, lower tunuarsuk member in profile 44c, kuugaartorfik, northern svartenhuk halvø. 251333 weakly plagioclase-phyric basalt enriched in incompatible trace elements (ba, th, u, nb, sr, ree). massive lava flow c. 60 m thick, middle tunuarsuk member in profile 37, nuup qaava, northern side of the innerit inlet on the west side of the innerit peninsula. 262815 aphyric basalt slightly enriched in incompatible trace elements (ba, th, u, nb, sr, ree, ti). massive lava flow c. 25 m thick, middle tunuarsuk member in profile 64, paannivik, southern innerit. 262793 plagioclase-phyric mg-poor basalt. lava flow 5-10 m thick, middle tunuarsuk member in profile 64, paannivik, southern innerit. 566511 aphyric basalt, nuuit member. highest lava flow in profile 65, salliaruseq, north-east of svartenhuk halvø. 566505 plagioclase microphyric basalt lava flow. reconnaissance sample, nunatak in side valley at the upper reaches of umiammakku sermia (glacier), gps coordinates 72°4.3′n, 51°58.8′w, altitude 1740 m (fig. 5), member status not certain, but note close similarity to sample 566511 from salliaruseq 75 km to the north-west. 263525 plagioclase-phyric basalt lava flow. middle part of nuuit member in profile 62, sigguk, western svartenhuk halvø. 181097 aphyric basalt lava flow. upper part of nuuit member in profile 40, pannertuup qaqqaa, east of the northernmost extension of ukkusissat fjord, north-east of svartenhuk halvø. 181102 plagioclase-glomerophyric basalt lava flow enriched in incompatible trace elements (ba, th, u, nb, sr, ree). lowest flow in nuuit member in profile 40, pannertuup qaqqaa, east of the northernmost extension of ukkusissat fjord, north-east of svartenhuk halvø. 566520 sparsely plagioclase-glomerophyric basalt lava flow, middle part of nuuit member in profile 65, salliaruseq, north-east of svartenhuk halvø. 262973 plagioclase-glomerophyric basalt lava flow. highest flow in nuuit member in profile 62, sigguk, western svartenhuk halvø. 251338 plagioclase-glomerophyric basalt lava flow enriched in incompatible trace elements (ba, th, u, nb, sr, ree). lowest flow in nuuit member in profile 37, nuup qaava, western innerit peninsula. 278520 weakly plagioclase-phyric magnesian basalt lava, lowest flow in skalø member in profile 50, umiiarfik south, northern svartenhuk halvø. 262896 plagioclase-phyric, 50 m thick basalt lava flow. lower skalø member in profile 41b, skalø south, southern skalø. 262754 plagioclase-phyric basalt lava flow. top flow in profile 64, paannivik, southern innerit peninsula. 262988 plagioclase-phyric basalt lava flow. middle skalø member in profile 62, sigguk, western svartenhuk halvø. 262912 plagioclase-phyric basalt lava flow. upper skalø member in profile 41b, skalø south, southern skalø. 262697 plagioclase-phyric basalt lava flow with very high tio2. skalø member, highest preserved flow in profile 55, aputituut, central svartenhuk halvø, http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 100 of 121 relatively evolved magmas produced by fractionation in large, presumably deep-seated magma chambers. the decrease in al2o3 and cao and increase in feo* in basalts with mg-number <57 reflect plagioclase fractionation (fig. 76). the stratigraphic variations in tio2 with height, as seen in fig. 42, reflect slow variations in the input/output ratio of magma in the chamber. the tunuarsuk member contains a significant component of picrites (with mg-numbers > c. 68), which the two younger members do not. this may be a result of decreasing magma production rates with time and thus longer residence times in the magma chambers. the very evolved lava flows on top of the skalø member on aputituut (profile 55) mentioned above may be interpreted as the products of prolonged fractionation in a dying magma chamber that was no longer replenished from deeper levels. concentrations of incompatible elements increase with increasing fractionation of the magma, and all three members follow similar fractionation trends (fig. 77). the significance of some relatively enriched samples and crossing and parallel ree and multi-element curves (fig. 79) is discussed in the later section on enriched mantle components. naqerloq formation arfertuarsuk member the arfertuarsuk member comprises tholeiitic basalts with relatively high contents of many incompatible elements. their levels of silica saturation (5–20 wt% normative hypersthene, average 13.0 wt%) are only slightly lower than for the basalts of the svartenhuk formation (5–23 wt% normative hypersthene, average 16.2 wt%). they are thus not alkaline, as alkali basalts by definition do not have normative hypersthene (le maitre 2002). most of them are not even transitional (i.e. with low normative hypersthene), although icelandic basalts with similar compositions are often called transitional or alkaline. major element variation diagrams for the arfertuarsuk member are shown in fig. 80 and ree and multi-element diagrams in fig. 81. other trace-element data are included in the variation diagrams for the svartenhuk formation in figs 77 and 78. the arfertuarsuk trachyte and its inclusions are compositionally far off scale in the variation diagrams but are shown in the ree and multi-element diagrams. representative chemical analyses are shown in table 7. the arfertuarsuk member basalts have 3.8–8.6 wt% mgo with an average of 5.8 wt% mgo. tio2 is in the range 1.8–4.0 wt%, and the average tio2 content is 3.0 wt% (79 analyses). the member is thus on average slightly more evolved to lower mgo and higher tio2 contents than the basalts of the svartenhuk formation, though there are large compositional overlaps for these elements (fig. 80). k2o and p2o5 contents are notably high, with averages of 0.80 wt% k2o and 0.39 wt% p2o5 compared to the svartenhuk formation with averages of 0.24 wt% k2o and 0.21 wt% p2o5; there are only small compositional overlaps for these elements (fig. 80). the p2o5/tio2 ratio is useful for distinguishing samples that have not been analysed for trace elements because this ratio is on average 0.10 ± 0.01 for the svartenhuk formation and 0.13 ± 0.01 for the arfertuarsuk member. the concentrations of other incompatible elements are also significantly higher in the basalts of the arfertuarsuk member than in those of the svartenhuk formation (fig. 77). moreover, many petrogenetically important element ratios such as p2o5/tio2, ba/sr, zr/y, nb/y, nb/zr, nb/la, la/sm and tb/lu are significantly increased relative to basalts of the svartenhuk formation. all arfertuarsuk member basalts have lan/smn >1 (see also figs 81, 83). the three uppermost flows in the arfertuarsuk profile show a reversal to less enriched compositions with element ratios moved towards those in the svartenhuk formation; these three flows are distinguishable in several diagrams, particularly in fig. 81. the basalt flows below and above the trachyte level are compositionally different; those below the trachyte level are less evolved than those above it. the former flows have mg-numbers of 55–58 and plot in tight clusters in many diagrams (figs 77, 80, 81). the arfertuarsuk trachyte flow has a very evolved composition and contains on average (5 analyses) 65.7 wt% sio2, 0.83 wt% tio2, 0.24 wt% mgo, 16.4 wt% al2o3, 6.1 wt% na2o, 5.2 wt% k2o and 0.26 wt% p₂o₅. it is anorthite normative and is thus not peralkaline; the normative content of albite + orthoclase is 85.5 %, reflecting the very high modal content of alkali feldspar. the trachyte magma had fractionated plagioclase, apatite and fe-ti oxide, as seen in the low concentrations of sr, p, and ti in the trace and multi-element diagrams (fig. 81). the composition of the syenite xenolith (table 7) shows that the xenoliths are closely related to the trachyte itself. they may represent fragments of deposits along the sidewall or in the top zone of the magma chamber where the trachyte melt fractionated. the dark, fine-grained inclusion (fig. 67) has the most extreme trace-element composition of all samples; its ree and multi-element patterns are parallel to those of the trachyte but at much higher levels, and relative losses of rb, ba, k, sr, p, eu and ti are evident. one possible explanation for this is that the inclusion represents the result of filter pressing of residual liquid into available pockets (a sort of aplite veins without quartz) that in the process left feldspar behind. this would increase the concentration of incompatible elements in the liquid and also explain the negative eu anomaly and the low rb, ba and k contents. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 101 of 121 40 45 50 55 60 30 35 40 45 50 55 60 65 70 75 80 85 mg -number svartenhuk formation sio 2 5 10 15 20 30 35 40 45 50 55 60 65 70 75 80 85 mg -number al2o3 5 10 15 20 30 35 40 45 50 55 60 65 70 75 80 85 mg -number feo* 0 5 10 15 30 35 40 45 50 55 60 65 70 75 80 85 mg -number cao 0 1 2 3 4 5 30 35 40 45 50 55 60 65 70 75 80 85 mg -number tio 2 0 1 2 3 4 5 30 35 40 45 50 55 60 65 70 75 80 85 mg -number na2o 0.0 0.5 1.0 1.5 2.0 30 35 40 45 50 55 60 65 70 75 80 85 mg -number k 2o 0.0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 30 35 40 45 50 55 60 65 70 75 80 85 mg -number p 2o5 arfertuarsuk member fig. 80 major-element variation diagrams for basalts of the naqerloq formation. a group of low-ti basalts from the lowest part of the formation below the arfertuarsuk trachyte flow is encircled in some diagrams. data for the svartenhuk formation in grey for comparison. vertical axes in wt% oxides. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 102 of 121 1 10 100 1000 la ce pr nd pm sm eu gd tb dy ho er tm yb lu naqerloq formation, arfertuarsuk member below trachyte level. trachyte and inclusions low-ti basalts trachyte fine -grained inclusion 1 10 100 1000 la ce pr nd pm sm eu gd tb dy ho er tm yb lu naqerloq formation, arfertuarsuk member above trachyte level 1 10 100 1000 rb ba th u nb ta k la ce pb pr sr nd p sm zr hf eu ti gd tb dy y yb low-ti basalts below trachyte flow. trachyte (two samples) fine -grained inclusion two intermediate inclusions 1 10 100 1000 rb ba th u nb ta k la ce pb pr sr nd p sm zr hf eu ti gd tb dy y yb three uppermost flows in arfertuarsuk profile fig. 81 ree and multi-element diagrams for representative rocks of the naqerloq formation. the different colours of the curves are intended to aid the distinction of the individual samples. left diagrams: chondrite normalised. right diagrams: primitive mantle normalised. normalisation factors from mcdonough & sun (1995). table 7 notes on analysed samples from the naqerloq formation and the delta-1 drill hole 262998 brown aphyric basalt relatively low in tio2. second lava flow in arfertuarsuk member, profile 62, sigguk, western svartenhuk halvø. 262879 brown aphyric basalt relatively low in tio2. first lava flow in arfertuarsuk member, profile 59, qooruusaq, western svartenhuk halvø. 278607 brown aphyric basalt, unlike other rocks in arfertuarsuk member not enriched in incompatible elements. one of three flows with this character, the three highest lava flows in profile 36b, arfertuarsuk, south-western svartenhuk halvø. 278601 brown aphyric basalt with an enriched composition typical of arfertuarsuk member. lava flow in the middle of profile 36b, arfertuarsuk, south-western svartenhuk halvø. 263512 strongly plagioclase-glomerophyric basalt. lava flow in the middle part of arfertuarsuk member in profile 62, sigguk, western svartenhuk halvø. 251372 plagioclase-clinopyroxene-phyric basalt. lava flow in the lower part of arfertuarsuk member in profile 41b, southern skalø. dated sample: 54.86 ± 0.44 ma (larsen et al. 2016). 263516 plagioclase-glomerophyric basalt. lava flow in the upper part of arfertuarsuk member in profile 62, sigguk, western svartenhuk halvø. 278588 plagioclase-glomerophyric basalt. lava flow in the middle part of arfertuarsuk member in profile 36b, arfertuarsuk, south-western svartenhuk halvø. 278660 anorthoclase trachyte lava flow, sample from the top of the flow. coastal exposure in innermost arfertuarsuk inlet, south-western svartenhuk halvø. 278669.2 decimetre-sized syenite xenolith in the arfertuarsuk trachyte flow near its base. coastal exposure in innermost arfertuarsuk inlet, southwestern svartenhuk halvø. 568601.2 dark, fine-grained xenolith in the arfertuarsuk trachyte flow (fig. 67). profile 67 in the upper qooruusaq valley, western svartenhuk halvø. swc-17 sidewall core at 2598 m depth in the delta-1 well. the sample is dated at 54.5 ± 1.5 ma. data from nelson et al. (2015). note the compositional similarity to the onshore naqerloq formation. the delta-1 well is situated c. 100 km wsw of svartenhuk halvø (fig. 1). http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 103 of 121 table 7 (continued) chemical analyses of rocks from the naqerloq formation and the delta-1 drill hole unit arfertuarsuk memer offshore ggu no. 262998 262879 278607 278601 263512 251372 263516 278588 278660 278669.2 568601.2 swc-17 profile no. 62 59 36b 36b 62 41b top 62 36b 67 well name sigguk qooruus. arfertuar. arfertuar. sigguk skalø sigguk arfertuar. arfertuar. arfertuar. qooruus. delta-1 major elements in wt% (xrf analyses) sio₂ 48.13 48.98 47.00 49.28 48.31 48.17 48.19 48.27 62.65 61.39 61.11 47.41 tio₂ 1.91 1.85 2.43 2.89 3.06 3.26 3.42 3.87 0.69 0.58 0.74 3.57 al₂o₃ 14.53 15.01 14.54 13.79 13.78 14.38 13.35 12.73 15.74 16.40 17.38 13.44 fe₂o₃ 3.88 4.48 5.17 6.65 6.40 5.42 5.74 5.62 2.22 2.66 8.17 15.35* feo 7.58 6.19 8.24 6.81 7.80 8.21 8.39 9.14 2.10 0.85 mno 0.19 0.16 0.21 0.19 0.26 0.19 0.22 0.22 0.15 0.49 0.12 0.22 mgo 7.36 6.80 6.46 5.60 5.17 4.86 5.15 4.66 0.50 0.27 0.08 4.83 cao 11.57 11.76 11.32 10.38 9.98 10.13 9.98 9.25 1.69 2.94 0.40 10.21 na₂o 2.42 2.39 2.28 2.70 2.79 2.77 2.82 2.77 5.82 5.87 6.36 2.67 k₂o 0.59 0.56 0.29 0.83 0.96 0.81 0.91 1.29 4.63 5.29 4.83 0.95 p₂o₅ 0.23 0.24 0.24 0.38 0.42 0.42 0.45 0.53 0.20 0.16 0.11 0.45 volatiles 1.64 1.56 2.10 1.16 1.21 1.32 1.42 1.30 1.74 1.82 0.46 0.29 sum 100.03 99.98 100.28 100.66 100.14 99.94 100.04 99.66 98.14 98.73 99.77 99.39 feo* 11.07 10.22 12.89 12.80 13.56 13.09 13.55 14.20 4.10 3.24 7.35 13.81 mg-no. 57.37 57.39 50.35 46.95 43.54 42.89 43.46 39.92 19.80 14.41 2.19 41.44 trace elements in ppm (icp-ms analyses) xrf anal. sc 35.9 37.98 40.0 42.8 32.38 32.96 33.68 41.2 6.19 5.58 5.54 27 v 308 312 364 321 324 347 359 350 13.4 6.14 24.9 336 cr 172 149 48.5 61.6 28.9 62.0 48.6 39.9 1.99 1.73 6.51 42 co 46.5 44.8 51.8 50.1 47.2 43.0 44.8 49.0 13.4 13.5 10.4 37 ni 92.3 83.8 64.5 57.7 41.4 45.9 44.1 44.1 1.44 1.55 11.3 46 cu 121 118 216 204 181 154 167 189 4.94 4.41 8.03 201 zn 82 85.3 108 122 128 130 135 150 62.7 55.6 306 123 ga 18.6 19.6 21.3 21.7 22.4 24.4 22.9 24.6 29.9 27.2 35.2 24 rb 10.1 9.26 2.71 18.9 18.3 12.1 14.3 33.0 56.3 42.5 24.1 9 sr 289 282 282 332 331 357 322 334 100 40.8 101 378 y 25.7 27.8 34.6 37.2 41.2 41.4 43.4 51.6 48.5 27.7 140 44.5 zr 103 109 157 219 248 244 258 321 293 167 271 275 nb 16.5 16.1 12.8 32.2 35.2 37.8 40.0 47.8 64.3 36.3 126 41.1 cs 0.030 0.118 0.041 0.091 0.056 0.124 0.092 0.259 0.108 0.107 0.084 ba 232 240 89.2 236 255 244 275 372 1303 684 164 348 la 13.3 13.8 11.5 23.3 25.4 25.0 27.0 35.2 48.7 26.9 132 ce 28.1 29.4 28.6 51.5 55.2 55.3 59.3 76.3 96.2 53.2 248 pr 3.83 4.01 4.33 6.97 7.51 7.83 8.38 10.2 11.9 6.64 31.3 nd 17.7 18.7 20.8 30.9 33.4 34.4 37.1 45.2 47.4 26.6 115 sm 4.40 4.60 5.77 7.38 8.16 8.25 8.74 10.6 9.70 5.43 21.7 eu 1.62 1.68 1.97 2.34 2.45 2.59 2.67 3.12 2.96 1.88 3.37 gd 4.74 4.89 6.29 7.92 8.49 8.76 9.12 11.2 9.15 5.20 21.2 tb 0.808 0.839 1.05 1.25 1.32 1.34 1.41 1.81 1.47 0.854 3.57 dy 4.52 4.90 6.13 7.01 7.45 7.61 8.05 9.65 8.44 4.77 21.4 ho 0.918 1.00 1.26 1.38 1.47 1.50 1.55 1.90 1.66 0.946 4.68 er 2.42 2.57 3.31 3.62 3.83 3.86 4.12 5.05 4.60 2.51 13.2 tm 0.349 0.388 0.458 0.505 0.566 0.569 0.574 0.702 0.682 0.372 1.94 yb 2.08 2.32 2.92 3.10 3.31 3.40 3.56 4.23 4.12 2.18 11.7 lu 0.317 0.336 0.432 0.457 0.477 0.486 0.494 0.597 0.614 0.352 1.84 hf 2.79 2.77 4.01 5.41 5.76 5.98 6.32 8.02 7.10 3.72 7.92 ta 0.957 0.980 0.876 2.23 2.08 2.15 2.33 3.24 3.17 1.84 8.30 pb 0.947 1.24 1.09 1.72 1.69 1.71 1.79 2.71 3.81 2.42 11.8 4 th 1.29 1.31 0.975 2.66 2.81 2.55 2.93 4.14 4.81 2.33 15.0 3 u 0.330 0.324 0.198 0.714 0.772 0.678 0.824 1.08 1.00 5.25 2.39 1 feo*= total iron calculated as feo. total iron determined as fe₂o₃ is marked by an * mg-number = 100 × atomic mg/(mg+fe²+), with the iron oxidation ratio adjusted to fe₂o₃/feo = 0.15. for notes on the samples, see previous page. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 104 of 121 intrusions the relatively few trace-element analyses of dykes and sills are plotted in figs 77, 78 and 82. representative chemical analyses are shown in table 8. a picrite dyke from the nunavik member and two strongly contaminated dykes similar to the kakilisaat member are easily identified chemically. the basaltic dykes form two compositional groups, one with a depleted and one with an enriched trace-element pattern, similar to the compositions of the svartenhuk formation and the naqerloq formation, respectively. presumably the dykes of the two groups fed the respective formations, and the compositional difference is therefore also an age indicator. this is supported by the results of chauvet et al. (2019) who presented 40ar–39ar ages for two dykes, one with ‘svartenhuk chemistry’ and an age of 58.98 ± 0.93 ma, and one with ‘naqerloq chemistry’ and an age of 54.41 ± 0.99 ma. the large sills along the cretaceous boundary fault system have compositions identical to the svartenhuk formation basalts (table 8). we therefore consider that these intrusions represent the svartenhuk formation and thus are of late paleocene age. this is in contrast to the sills intruded in a similar setting along the boundary fault on nuussuaq, which are of eocene age and have corresponding enriched compositions (storey et al. 1998; larsen et al. 2009). primary magmas, melting conditions and development with time vaigat formation as described above, the earliest volcanic unit deposited in eastern svartenhuk halvø, the kakilisaat member of the vaigat formation, is thoroughly crustally contaminated and also bears an imprint of an enriched component. the following unit, the nerutusoq member, is much less contaminated and contains a distinct enriched component. we consider that these two members are related and ultimately derived from similar parental magmas generated in the asthenosphere. crustal contamination and trace-element enrichment are post-generation modifications, and consequently the asthenosphere-derived primary magma for the kakilisaat and nerutusoq members cannot be characterised in detail. it was, however, picritic because the high cr contents (>1000 ppm) in the least contaminated samples indicate derivation from magmas with ≥15 wt% mgo (fig. 74). the simplest assumption is that the primary magma was similar to that of the overlying nunavik member. there is no evidence of smaller degrees of melting than in the nunavik member as suggested by agranier et al. (2019). the post-melting modifications were achieved when the magmas intruded the local lithosphere and came into contact with crustal material and old enriched material. after this stage, the magmas of the nunavik member were able to pass unchanged through the lithosphere to the surface without stalling in magma chambers and becoming contaminated on the way. the unmodified parental magma of the nunavik member was neither crustally contaminated nor did it contain a very enriched component. the strong compositional similarity to the ordlingassoq member of the vaigat formation on nuussuaq and disko includes the slope of both the light ree and the heavy ree patterns, with tbn/lun averaging 1.61 ± 0.15 for the nunavik member and 1.63 ± 0.15 for the ordlingassoq member (pedersen et al. 2018), indicating residual garnet in the source mantle (fig. 83). the sr-nd isotope ratios are also similar, with both members mostly having 87sr/86sr = 0.7031–0.7034 and 143nd/144nd = 0.5129–0.5130 (holm et al. 1993; lightfoot et al. 1997; larsen & pedersen 2009; agranier et al. 2019). the melting conditions for 1 10 100 1000 la ce pr nd pm sm eu gd tb dy ho er tm yb lu naqerloq fm dyke svartenhuk fm dyke kakilisaat mb dyke nunavik mb dyke dykes 1 10 100 1000 rb ba th u nb ta k la ce pb pr sr nd p sm zr hf eu ti gd tb dy y yb fig. 82 ree and multi-element diagrams for dykes. the distinction between dykes of the svartenhuk and naqerloq formations is described in the text; see also fig. 85. the ta peak in the multi-element pattern of the nunavik member picrite dyke is a slight contamination from the grinding vessel. left diagram: chondrite normalised. right diagram: primitive mantle normalised. normalisation factors from mcdonough & sun (1995). http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 105 of 121 table 8 chemical analyses of dykes, sills and samples dredged from baffin bay unit dykes sills dredged samples ggu no. 251511 251387 165242 262982 262799 262994 263522 278399 278320 251385 bb8-13d8 bb8-12d2 profile no. 9 2 31 62 64b 62 62 52 44c upernavik escarpment location pk 1078 qinngusa. aputit.qaq. sigguk paannivik sigguk sigguk nuuit qaq. simiutaq qinngusa. baffin bay baffin bay affinity kakilis. mb nunav. mb sva. fm sva. fm sva. fm naq. fm naq. fm naq. fm sva. fm sva. fm sva. fm sva. fm sio₂ 49.63 46.11 48.32 49.48 48.59 45.18 46.85 48.60 47.86 48.86 47.77 48.21 tio₂ 1.36 1.16 1.72 2.32 3.22 3.41 4.03 3.49 1.88 2.26 2.77 2.53 al₂o₃ 13.67 12.08 13.65 14.12 13.39 14.16 12.60 13.60 15.56 14.83 13.28 13.20 fe₂o₃ 2.26 3.52 3.25 4.80 3.74 2.69 5.39 6.25 3.80 5.12 16.01* 16.22* feo 7.94 7.48 8.15 8.50 10.73 10.87 10.03 6.79 7.68 8.09 mno 0.16 0.17 0.22 0.23 0.23 0.22 0.23 0.20 0.17 0.20 0.251 0.191 mgo 9.69 16.24 9.21 5.96 5.45 6.57 4.97 4.15 6.46 5.18 6.16 5.21 cao 10.41 9.69 11.69 11.02 10.22 11.66 9.93 9.14 11.96 10.68 10.35 9.58 na₂o 1.99 1.63 2.02 2.68 2.51 2.39 2.93 3.11 2.27 2.64 2.67 2.90 k₂o 0.19 0.09 0.17 0.40 0.45 0.93 1.20 1.25 0.19 0.33 0.25 0.53 p₂o₅ 0.14 0.12 0.15 0.23 0.37 0.43 0.52 0.48 0.19 0.23 0.27 0.30 volatiles 2.32 1.82 1.07 0.84 1.23 1.45 1.30 2.61 2.24 1.57 sum 99.76 100.11 99.62 100.57 100.12 99.95 99.98 99.67 100.26 99.99 99.78 98.87 feo* 9.97 10.65 11.07 12.82 14.09 13.29 14.88 12.41 11.10 12.70 14.41 14.59 mg-no 66.28 75.53 62.71 48.47 43.89 50.00 40.32 40.34 54.07 45.21 46.39 41.94 trace elements in ppm (icp-ms analyses) sc 36.2 34.5 39.8 40.6 41.0 30.7 33.2 29.7 36.1 36.9 45 42 v 272 290 356 435 511 384 399 382 365 398 487 416 cr 607 1202 518 99.4 86.5 61.9 30.5 13.5 234 46.8 89 45 co 52.1 68.8 54.1 47.1 45.8 53.4 48.5 79.6 44.8 40.6 60 57 ni 257 621 222 63.0 69.0 81.1 38.6 94.6 102 45.7 68 52 cu 104 110 169 286 449 197 271 195 188 247 291 301 zn 77.3 79.6 83.7 110 132 113 151 150 91.9 104 127 121 ga 18.9 17.6 18.4 21.2 22.9 20.5 23.9 24.4 21.4 21.9 rb 5.56 2.12 2.89 10.1 14.9 24.0 27.1 31.6 2.68 6.25 2.0 7.5 sr 183 121 180 203 194 559 324 332 257 211 203 209 y 24.5 19.3 27.1 38.5 56.7 28.9 48.6 48.3 27.7 36.7 46 46 zr 107 59.8 91.0 150 229 185 273 306 115 136 163 164 nb 6.75 2.24 6.29 9.33 14.1 35.8 43.3 47.9 6.72 7.82 11.4 13.0 cs 0.247 0.088 0.059 0.286 0.217 0.203 0.158 0.175 0.058 0.064 0.04 0.06 ba 116 14.22 43.6 64.4 89.0 361 330 346 50.5 67.3 73 88 la 10.1 2.46 7.24 9.68 14.8 27.0 30.1 32.6 7.15 9.53 14.1 13.4 ce 22.9 7.36 17.3 23.8 36.1 57.9 64.4 71.7 19.2 23.4 35.7 33.5 pr 3.20 1.32 2.53 3.80 5.48 7.55 8.79 9.89 2.99 3.61 4.9 4.7 nd 14.2 7.30 12.3 18.7 25.9 33.0 39.0 42.6 15.1 17.2 24.1 24.3 sm 3.77 2.44 3.61 5.46 7.63 7.04 9.20 9.92 4.43 5.22 6.68 6.98 eu 1.19 0.92 1.30 1.86 2.34 2.31 2.86 2.94 1.55 1.81 2.41 2.57 gd 4.30 2.99 4.60 6.41 9.14 6.80 9.92 10.1 5.21 6.46 6.43 6.39 tb 0.714 0.513 0.772 1.11 1.53 1.02 1.52 1.55 0.850 1.08 1.31 1.31 dy 4.22 3.27 4.88 6.50 9.14 5.65 8.94 8.59 5.14 6.76 7.58 7.57 ho 0.885 0.671 1.01 1.39 1.95 1.06 1.73 1.68 1.03 1.38 1.59 1.57 er 2.36 1.84 2.71 3.75 5.20 2.70 4.55 4.35 2.73 3.80 4.53 4.39 tm 0.356 0.276 0.402 0.572 0.797 0.363 0.649 0.627 0.389 0.545 0.63 0.58 yb 2.10 1.70 2.44 3.41 4.81 2.16 3.87 3.72 2.38 3.36 3.69 3.61 lu 0.325 0.247 0.373 0.520 0.713 0.325 0.575 0.553 0.341 0.517 0.44 0.44 hf 2.78 1.62 2.60 4.12 5.77 4.58 6.67 7.16 3.08 3.67 4.18 4.40 ta 0.455 0.219 0.514 0.540 0.867 2.09 2.58 2.85 0.453 0.514 0.65 0.75 pb 2.55 0.323 0.695 1.03 1.70 1.73 2.20 1.30 0.846 1.58 3 2 th 1.64 0.165 0.479 0.824 1.54 2.47 3.27 3.76 0.546 0.946 2.26 1.34 u 0.384 0.048 0.147 0.269 0.515 0.696 1.12 1.06 0.176 0.273 0.49 0.39 feo*= total iron calculated as feo. total iron determined as fe₂o₃ is marked by an * mg-number = 100 × atomic mg/(mg+fe²+), with the iron oxidation ratio adjusted to fe₂o₃/feo = 0.15. for notes on the samples, see next page. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 106 of 121 the primary magmas of the vaigat formation modelled by larsen & pedersen (2009) can thus be directly applied to the svartenhuk halvø area: the melts were formed beneath a lithospheric lid 80–100 km thick by 15–20% melting of an asthenospheric mantle with residual garnet in the source. the mantle was probably heterogeneous, with highly depleted parts giving rise to magmas with lan/smn <1 (fig. 83) and zr/y and nb/y ratios below the iceland field (fig. 84). less-depleted parts (the ‘icelandic’ component of holm et al. 1993) gave rise to magmas with lan/smn >1 (fig. 83) and zr/y and nb/y ratios inside the iceland field (fig. 84), all within the same upwelling melting column. the resulting melts were picritic with about 17.5 wt% mgo and passed quickly through the lithosphere to the surface. svartenhuk formation the svartenhuk formation is 1–2 million years younger than the vaigat formation (fig. 3). the change from dominantly picritic magmas of the vaigat formation to dominantly basaltic magmas of the svartenhuk formation must have involved establishment of large, deep magma chambers, probably at the mantle/crust boundary, in which the magmas fractionated. the same change took place between the vaigat formation and the maligât formation on nuussuaq and disko (larsen & pedersen 2009). considering the age difference between the maligât and svartenhuk formations (fig. 3), this change was connected with the demise of the vaigat formation magmatism and could have been caused by changed tectonic conditions. conditions for magma generation in the asthenosphere stayed much the same, as evidenced by key incompatible element ratios. the light and heavy ree ratios for the svartenhuk formation plot mainly in the same area as the nunavik member (fig. 83), suggesting similar melting conditions and that the primary magmas were still picritic. samples with relatively high lan/ smn ratios are discussed below. the zr/y vs. nb/y relations likewise suggest a mantle source similar to that of the nunavik member, with compositions plotting both within and below the iceland field (fig. 84). sr-nd isotope data for the svartenhuk formation are sparse, but the available results suggest isotope ratios similar to those of the nunavik member (o’nions & clarke 1972; holm et al. 1993; agranier et al. 2019). a development with time is seen in decreasing heavy ree ratios. whereas the nunavik member has an average tbn/lun of 1.61 ± 0.15 (enriched samples excluded), tbn/lun averages decrease from 1.56 ± 0.15 in the tunuarsuk member to 1.50 ± 0.18 in the nuuit member and 1.45 ± 0.08 in the skalø member (fig. 83). as the mantle source was apparently similar to the source of the nunavik member, the temporal development in table 8 (continued) notes on analysed samples of dykes and sills and samples dredged from baffin bay 251511 silicic magnesian basalt, 40 cm wide dyke in hyaloclastite of the kakilisaat member of the vaigat formation, probably a feeder dyke for this member. profile 9, peak 1078 m, eastern svartenhuk halvø. 251387 olivine-phyric picrite, 3 m wide dyke in picrite lava flows of the nunavik member of the vaigat formation, probably a feeder dyke for this member. qinngusaaq mountain, eastern svartenhuk halvø east of the cretaceous boundary fault system. 165242 olivine-phyric basalt, 5–6 m wide dyke in lava flows of the skalø member of the svartenhuk formation. geochemical affinity to the svartenhuk formation. profile 31, aputituut qaqqaat, south-central svartenhuk halvø. 262982 plagioclase-glomerophyric basalt, 2 m wide, coast-parallel dyke in lava flows of the skalø member of the svartenhuk formation. geochemical affinity to the svartenhuk formation. profile 62, sigguk, western svartenhuk halvø. 262799 basalt dyke in lava flows of the tunuarsuk member of the svartenhuk formation. geochemical affinity to the svartenhuk formation. profile 64, paannivik, southern innerit peninsula. 262994 aphyric basalt, 1.2 m wide, nw-striking dyke in lava flows of the skalø member of the svartenhuk formation. geochemical affinity to the naqerloq formation. profile 62, sigguk, western svartenhuk halvø. 263522 aphyric basalt dyke in lava flows of the nuuit member of the svartenhuk formation. geochemical affinity to the naqerloq formation. coastal cliff in profile 62, sigguk, western svartenhuk halvø. 278399 aphyric basalt, 2.5 m wide dyke in fault zone in lava flows of the skalø member of the svartenhuk formation. geochemical affinity to the naqerloq formation. profile 52, nuuit qaqqaat, north-western svartenhuk halvø. 278320 basalt (dolerite), c. 50 m thick transgressive sill in lava flows of the nuuit member of the svartenhuk formation close to the cretaceous boundary fault system (fig. 6). geochemical affinity to the svartenhuk formation. profile 44c, kuugaartorfik, northern svartenhk halvø. 251385 plagioclase-glomerophyric basalt, up to 30 m thick sill in contact with gneiss and cutting pillow breccias and lava flows of the vaigat formation (fig. 6). geochemical affinity to the svartenhuk formation. qinngusaaq mountain, eastern svartenhuk halvø east of the cretaceous boundary fault system. bb8-13d8 basalt dredged on the upernavik escarpment in baffin bay at 73°05.5′ n, 58°09.1′ w, c. 80 km north-west of upernavik. full sample number bbs08-13d-8. data from polteau & planke 2008. note geochemical affinity to the svartenhuk formation. bb8-12d2 basalt dredged on the upernavik escarpment in baffin bay at 73°06.6′ n, 58°16.2′ w, c. 80 km north-west of upernavik. full sample number bbs08-12d-2. data from polteau & planke 2008. note geochemical affinity to the svartenhuk formation. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 107 of 121 tbn/lun must have been caused by melting dynamics and is interpreted to indicate a thinning lithospheric lid and shallowing melting in the garnet-to-spinel transition interval at 100–80 km depth, with decreasing amounts of garnet in the source. the basalts of the svartenhuk formation show small but significant compositional differences from the uncontaminated basalts of the maligât formation (the rinks dal member) on disko and nuussuaq. whereas in practice the major elements are indistinguishable, trace-element ratios (figs 83 and 84) show that the svartenhuk formation basalts are shifted to higher lan/smn and higher nb/y than most of the maligât formation basalts, although there is a partial overlap with the maligât formation. the svartenhuk formation basalts have lan/smn in the range 0.8–1.6 and there0.0 0.5 1.0 1.5 2.0 2.5 3.0 0 1 2 3 4 tb n /l u n lan/smn nunavik mb vaigat formation ordlingassoq mb field 20 10 15 10 15 dm pm a 0.0 0.5 1.0 1.5 2.0 2.5 3.0 0 1 2 3 4 tb n /l u n lan/smn tunuarsuk mb nuuit mb skalø mb arfertuarsuk mb dykes & sills naqerloq formation svartenhuk formation b maligât fm rinks dal mb field 0.01 0.10 1.00 10.00 1 10 n b/ y zr/y nunavik mb nerutusoq mb kakilisaat mb vaigat fm ordlingassoq mb 0.01 0.10 1.00 10.00 1 10 n b/ y zr/y tunuarsuk mb nuuit mb skalø mb arfertuarsuk mb dykes & sills maligât fm rinks dal mb fig. 83 ree cross plots and mantle melting models for the primary melts of the volcanic rocks of the svartenhuk region; enriched and crustally contaminated samples are not plotted. data are chondrite normalised (n). a: nunavik member of the vaigat formation. the black contour line is the field of the ordlingassoq member of the vaigat formation on nuussuaq and disko, with which the nunavik member is correlated. red curves for melting in garnet facies are shown for depleted mantle (dm) and primitive mantle (pm). numbers on the melting curves indicate degrees of melting (%). same melting models as in larsen & pedersen (2009). mantle modes and melting modes are from mckenzie & o’nions (1991). melting type is non-modal batch melting. dm trace-element starting composition is from mckenzie & o’nions (1991). pm starting composition is from mcdonough & sun (1995). partition coefficients are from mckenzie & o’nions (1991). b: svartenhuk and naqerloq formations. the black contour line is the field of the rinks dal member of the maligât formation on nuussuaq and disko; note the difference in lan/smn between the svartenhuk and maligât formations and their similar values of tbn/lun. fig. 84 nb/y vs. zr/y diagrams. a: vaigat formation. black contour line is the field of the ordlingassoq member on nuussuaq and disko; note the exact resemblance to the nunavik member. b: svartenhuk and naqerloq formations. black contour line is the field of the rinks dal member of the maligât formation on nuussuaq and disko; note the difference to the svartenhuk formation. the two parallel lines are the upper and lower bounds of the empirical field of data from iceland, after fitton et al. (1997). rocks more depleted in nb, such as many mid-ocean-ridge basalts, plot below the iceland field. this diagram is useful for distinguishing between ‘icelandic-type’ mantle sources and more depleted sources (e.g. saunders et al. 1997; fitton et al. 1997, 1998; larsen & pedersen 2009). http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 108 of 121 fore the light ree curves cross each other, as illustrated in fig. 79. in contrast almost all maligât formation basalts have lan/smn <1 and near-parallel ree curves (larsen & pedersen 2009; pedersen et al. 2018). the two formations are best discriminated in a plot of ce/y vs. zr/nb (fig. 85) where they occupy separate fields for zr/nb <19; the overlap area with zr/nb >19 is only sparsely occupied by svartenhuk formation basalts. the naqerloq formation occupies a separate field. based on this diagram, larsen et al. (2016) and pedersen et al. (2018, fig. 101) found that the basalt dykes on disko and nuussuaq are different from the basalt lava flows of the maligât formation but similar to the lava flows of the svartenhuk and naqerloq formations. the dyke ages yield similar results (fig. 3). the basalt dykes on disko and nuussuaq are therefore thought to have fed lava flows of the svartenhuk and naqerloq formations, which must have been present in that area but are now removed by erosion. a possible interpretation of fig. 85 is that the svartenhuk formation basalts are mixed with an enriched component that is not present in the maligât formation basalts. naqerloq formation the naqerloq formation is 2–3 million years younger than the svartenhuk formation and the basalts of the two formations are significantly different. as described above, the naqerloq formation basalts have higher contents of incompatible elements (figs 77, 80), higher zr/y and nb/y ratios plotting within the iceland field (fig. 84), higher la/sm and tb/lu ratios (fig. 83) with a tbn/lun average of 1.85 ± 0.12, higher ce/y and lower zr/nb (fig. 85). the 87sr/86sr ratios are slightly higher, around 0.7036, whereas the 143nd/144nd ratios are similar, 0.5129–0.5130 (holm et al. 1992, p. 352 (kanísut member); agranier et al. 2019). the naqerloq formation melts must have been generated either by smaller degrees of melting of a mantle source similar to the earlier source, or by melting of an enriched mantle source with higher contents of trace elements than primitive mantle (fig. 83). smaller degrees of melting of a mantle source similar to the earlier one should produce differences in the major elements, for example higher na2o, which are not seen (fig. 80), and the small differences in the sr isotope ratios also favour a different mantle source. the melting modelling in fig. 83 is based on depleted and primitive mantle compositions, and therefore the degrees and depths of melting of the primary magmas of the naqerloq formation cannot be modelled. it is possible that the degrees of melting were smaller than for the magmas of the svartenhuk formation, as suggested by agranier et al. (2019), but this is not necessary to explain the data. storey et al. (1998) suggested that the ‘new’ mantle material from which the naqerloq formation was derived, was a remnant of the iceland plume head left stranded beneath the west greenland lithosphere. however, deposition of the naqerloq formation was not a local phenomenon, and volcanic rocks chemically similar to those of the naqerloq formation are found at several places in and around greenland. they were drilled in the delta-1 well offshore svartenhuk halvø (nelson et al. 2015) and dredged in the labrador sea south of the hecla high volcanic centre, which was interpreted as their most probable source (fig. 1; larsen 0.0 0.5 1.0 1.5 2.0 0 5 10 15 20 25 30 c e/ y zr/nb maligât formation svartenhuk formation enriched svartenhuk fm naqerloq formation dykes & sills fig. 85 ce/y vs. zr/nb diagram demonstrating the compositional differences between the svartenhuk formation and the rinks dal member of the maligât formation on disko and nuussuaq. there is an area of overlap at high zr/ nb, but for zr/nb <19 the separation of the two formations is good. the enriched svartenhuk formation basalts form a field extending towards the naqerloq formation, which occupies a separate field. the intrusions form two groups corresponding to the svartenhuk and naqerloq formations. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 109 of 121 & dalhoff 2006). ocean drilling program (odp) hole 918 offshore south-east greenland contains a sill with a similar composition (fitton et al. 1998), and the flood basalts in the scoresbysund area include similar units, particularly the rømer fjord formation (larsen et al. 1989). as far as these rocks have been dated they are of earliest eocene age, 56–50 ma. thus the ‘new’ mantle can rather be interpreted as a major part of the iceland plume that spread through the north atlantic during the phase of plate rearrangement and change of spreading direction from ne–sw to nearly n–s, that took place at the transition from paleocene to eocene (oakey & chalmers 2012). enriched mantle components basalts with a chemical imprint of material enriched in the most incompatible trace elements occur in all members of the vaigat and svartenhuk formations and are most common in the tunuarsuk member. the patterns of enrichment are variable, as illustrated by the multi-element patterns for selected samples in fig. 86. compared to the normal basalts with nbn/lan ~ 1 , some samples have nbn/lan <1, others have nbn/lan >1. some samples have ba peaks, one has a u peak, the nunavik picrite has a distinct nb–ta peak, some samples have no extra pb and ti shows both peaks and troughs. all samples have increased light ree (la–sm), and many except the enriched nuuit member samples have steeper heavy ree limbs (gd–yb) and even lower y and yb than the normal basalt. because of the variability it has not been possible to make a rigid definition of ‘enriched’ samples; they are generally characterised by having high nb and high la–nd relative to their degree of fractionation. three enriched samples from the nunavik and tunuarsuk members were analysed for sr and nd isotopes by holm et al. (1993); they have high 87sr/86sr ratios (0.7036–0.7041) and low 143nd/144nd ratios (0.5128– 0.5129) relative to the normal basalts and picrites, indicating a separate source of the enriched material. sporadic addition of enriched lithospheric material in the nunavik, tunuarsuk and skalø members chemically enriched lavas occur as individual flows surrounded by flows that are chemically normal. we consider it most probable that the magmas of these flows have acquired their enrichment on their way to the surface after they left the melting column (nunavik member, the βo1 marker unit) and the deep-seated magma chambers (tunuarsuk and skalø members). the enriched components therefore resided in the lithospheric mantle or crust and represent veins or dykes of older alkaline rocks, which could easily be partially melted and assimilated by batches of passing hot basaltic and picritic melts on their way to the surface. the resulting enrichment patterns would depend on the melting mineral assemblage. older alkaline carbonate-rich rocks are reported from the region, e.g. from the islands of qinngusaaq and qeqertarsuaq, both c. 30 km east of southern svartenhuk halvø (henderson & pulvertaft 1987; knudsen et al. 2010; mott et al. 2013). a similar origin has been proposed for variously enriched rocks akin to those on svartenhuk halvø from the manîtdlat member in the vaigat formation on disko, some 150 km farther south (larsen et al. 2003; this minor member is associated with the ordlingassoq member but not listed in fig. 3). in detail, however, the occurrences are different; the rocks of the manîtdlat member are mostly more enriched than those on 0 1 10 100 rb ba th u nb ta k la ce pb pr sr nd p sm zr hf eu ti gd tb dy y yb s am pl e/ p rim iti ve m an tle skalø member nunavik member nuuit member nerutusoq member tunuarsuk member normal svartenhuk fm variously enriched samples fig. 86 multi-element diagram for selected samples showing variable patterns of element enrichment. ratios between normalised elements such as nb/la and th/u vary from greater than one to less than one. two samples show no pb enrichment, ti shows troughs, plains or peaks, and the steepness of the gd–yb limb is variable. samples as as follows: nunavik member: 251527 (βo1 flow); nerutusoq member: 181013 and 263823; normal svartenhuk formation: 566518; tunuarsuk member: 251333 and 262815; nuuit member: 251338; skalø member: 278622. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 110 of 121 svartenhuk halvø, and their sr-nd isotope compositions are dissimilar to the few results published by holm et al. (1993). an enriched asthenospheric mantle component the nuuit member shows a systematic up-section variation of the most incompatible trace elements in most analysed profiles (fig. 87). the basalts identified as enriched all occur at the base of the member, but not all basal samples are classified as enriched. the lowest samples, whether they are enriched or normal, have relatively high lan/ndn, nb/zr, ba/sr and other incompatible-element ratios, and with few exceptions the ratios decline systematically up-section. the variation is most clearly seen in the lan/ndn ratio (fig. 87). this development is seen in six out of eight profiles analysed systematically for trace elements. in contrast, the heavy ree ratios (tbn/lun) of all samples are near-constant, suggesting melt generation at near-similar depths. there is no correlation between lan/ndn and the degree of fractionation of the erupted magma as indicated by the mgo and tio2 contents (compare fig. 87 with the tio2 profiles in fig. 42). if the cause of the variation is increasing degrees of melting up-section there should also be differences in other elements, which are not seen. the simplest explanation of the data is that the melting asthenospheric mantle comprised two components, one relatively depleted in incompatible trace elements and another relatively enriched in such elements. melt from the enriched component was mainly produced at the start of the nuuit member and gradually diminished during continued melting. the two profiles closest to the coast (sigguk and fladø) do not show the same pattern as the others, which suggests separate magma chambers beneath the coastal and inland areas. 0.6 0.8 1.0 1.2 1.4 1.6 1.8 2.0 w sw :c oa st al z on e in la nd p ro fil es e n e : lan/ndn eastern nunatak salliaruseq pannertuup qaqqaa qaqqap qaa marrarnaq nuup qaava sigguk fladø stratigraphic up in all pro�les 0.6 0.8 1.0 1.2 1.4 1.6 1.8 2.0 s tra tig ra ph ic u p lan/ndn paannivik skalø member nuuit member tunuarsuk member b a fig. 87 stratigraphic variation up-section of lan/ndn (chondrite normalised) in eight densely sampled and analysed profiles. a: data for the nuuit member in seven profiles stacked vertically in a wsw–ene succession. the lowest profiles are those within the coastal zone and the highest profiles are those farthest inland. different colours and shapes are for clarity. b: up-section variation through the whole svartenhuk formation in the paannivik profile. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 111 of 121 melting and mixing relations in the svartenhuk formation only one analysed sample profile covers all three volcanic members of the svartenhuk formation: the paannivik profile (64). the plot of lan/ndn vs. stratigraphic height for this profile (fig. 87b) shows that the three members have different styles of up-section variation. the tunuarsuk member shows a large compositional spread and no systematic up-section variation; the nuuit member shows the systematic decrease in lan/ndn described above, and the skalø member has constant lan/ndn throughout. other analysed profiles through the tunuarsuk and skalø members are less complete but indicate the same pattern of overall variation as at paannivik. the data may be interpreted in terms of melting and mixing dynamics. in the tunuarsuk member, the melt batches from the mantle did not mix efficiently in the magma chamber where they fractionated, and both very depleted melts with lan/ndn around 0.8 and very enriched melts with lan/ndn up to 1.6 were preserved and erupted. this may have been caused by relatively short residence times in the magma chamber, which also led to more magnesian basalts than in the two subsequent members. in the nuuit member mixing of the individual melt batches intruded into the magma chamber became much more efficient, though not complete in the beginning, and at the same time the amount of enriched melt gradually diminished with further melting. in the skalø member, mixing was either perfect or the enriched mantle component had been exhausted; the low lan/ndn ratio of 0.9 suggests the latter. our interpretations can be tested with isotope data, which are, however, currently not available. a corollary of the interpretations is that the enriched flows in the tunuarsuk member may have acquired their enrichment in two different ways, in the asthenosphere and in the lithosphere, which would contribute to their observed compositional variability. the existence of one or more enriched components in the basalts is precisely the reason why basalts of the svartenhuk formation can be chemically distinguished from basalts of the maligât formation, which lack such a component. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 112 of 121 concluding remarks the volcanic succession in the northern part of the nuussuaq basin, on svartenhuk halvø and the areas north and north-east of this, was deposited during a series of geological events that affected the entire basin. the northern succession is therefore in general correlatable to the more southerly parts, but on the formation and member level there are also distinct differences in the geological evolution. volcanism on svartenhuk halvø may have started somewhat later than in the southern nuussuaq basin. the nunavik member correlates excellently with the upper part of the vaigat formation (ordlingassoq member) in the south, but older members of the vaigat formation exposed on disko and nuussuaq may be present at depth beneath the western svartenhuk halvø. at the end of the volcanic activity of the vaigat formation, the mg-rich volcanic rocks formed a n–s elongated, mainly subaerially exposed dome that extended for c. 250 km from central disko and nuussuaq over ubekendt ejland, where thicknesses appear to be greatest, and across svartenhuk halvø to the southern innerit peninsula, where the formation thins and tapers out. after this stage, the conditions that had allowed picritic magmas to pass relatively unmodified to the surface in the whole nuussuaq basin changed. deep magma chambers developed, in which the picritic magmas evolved to basaltic compositions, and the modes of evolution in the north and south diverged. volcanism in the south continued with the extrusion of the basalts of the maligât formation, whereas in the north there was a hiatus during which quartzofeldspathic sediments were deposited. when volcanism resumed in the north, the basalts of the svartenhuk formation were deposited; the bulk of these are younger than the maligât formation. the three members of the svartenhuk formation are mappable because of morphological and colour differences. they show a large compositional overlap, but up-section variations produced by processes in the magma chambers support the individual nature of the members. the svartenhuk formation is probably extensively present offshore where it has been recognised in drill hole alpha-1 and in dredged samples. onshore the formation is also present on ubekendt ejland and probably in western nuussuaq and as dykes on disko and nuussuaq. at the transition from paleocene to eocene time, there was a volcanic hiatus that lasted for c. 1–2 ma during which tectonic movements and reorganisation of plate spreading directions took place. volcanism resumed with extrusion of the naqerloq formation basalts, which are compositionally different from the paleocene basalts and were generated in a different mantle. lava successions of the naqerloq formation are preserved only in the westernmost parts of the nuussuaq basin from hareøen to svartenhuk halvø and mainly as erosional remnants, but dykes throughout disko, nuussuaq and svartenhuk halvø show that the formation was erupted over large parts of the nuussuaq basin. it is probably also extensively present offshore where it has been recognised in the drill hole delta-1. further studies in the northern part of the nuussuaq basin may provide a better understanding of some aspects of the geology. first, the volcanic succession east of c. 54°w is poorly known. our present knowledge is based on extensive photogrammetric studies whereas ground studies only comprise two sampled profiles. second, formation boundaries with sediments should be systematically investigated to provide a more precise understanding of the events that took place at the time; this is particularly true for the paleocene–eocene transition. third, the uncertainties concerning the internal stratigraphy of the naqerloq formation in the western svartenhuk halvø area may be resolved with more detailed work. fourth, the northernmost area (north of 72°30′n) remains completely unvisited. acknowledgements over the long span of years, we have been indebted to many persons. in the field, karl aage jørgensen and troels nielsen made substantial and invaluable contributions to sampling and logging of basalt profiles and frank andreasen provided information on interbasaltic sediments. we also owe our thanks to expedition leader gilroy henderson and skipper andreas viðstein on m/s steenstrup, as well as to vice director flemming getreuer christiansen and department leaders niels henriksen, christian knudsen and stefan bernstein for general ggu/geus support. hans jepsen provided guidance in the use of the kern pg2 stereoplotter in ggu’s photogeological laboratory. jørgen kystol and ib sørensen and in later years olga nielsen at ggu/geus’s rock geochemical laboratory maintained the constantly high quality of the xrf and icp-ms analyses of the rocks. we are particularly grateful to the late chris pulvertaft for his pioneering field work in the region in the 1960s and for his unswerving support and interest in svartenhuk halvø. we extend our thanks to erik vest sørensen, kristian svennevig and michelle de wolfe for their recent sampling of some key localities and discussions of the field relations. the two first-named also generously put their more recent photographs at our disposal, as did asger ken pedersen. henrik nøhr-hansen carried out the micropalaeontological analyses of sediment samples. jette halskov patiently and expertly prepared the drawings and annotated the photographs. the pen drawings in figs 14–18 were digitised and further adjusted by ane asmussen. jacob bendtsen helped with the photoshop work. the digital elevation map in fig. 9 was constructed by christian brogaard pedersen. an early draft of the manuscript and some figures were carefully commented and amended by stuart watt. special thanks go to asger ken pedersen for numerous discussions throughout the work and for comments on the manuscript. we are further grateful to the reviewers david peate and richard wilson for their helpful and constructive reviews. jgl extends his thanks to his family for their year-long patience. http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 113 of 121 additional information funding statement the work was primarily funded by ggu and later geus. the report by larsen & larsen (2010) was funded by a consortium of oil companies. competing interests the authors declare no competing interests. author contributions jgl: field work, map compilation, writing original draft, revision. lml: writing revision of original draft, new chapter on geochemistry and petrology, figures, revision, editing. additional files four supplementary files are available at: https://doi.org/10.22008/fk2/ef2nkr http://www.geusbulletin.org https://doi.org/10.22008/fk2/ef2nkr www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 114 of 121 references abdelmalak, m.m., geoffroy, l., angelier, j., bonin, b., callot, j.p., gélard, j.p. & aubourg, c. 2012: stress fields acting during lithosphere breakup above a melting mantle: a case example in west greenland. tectonophysics 581, 132–143. https://doi.org/10.1016/j. tecto.2011.11.020 agranier, a., maury, r.c., geoffroy, l., chauvet, f., le gall, b. & viana, a.r. 2019: volcanic record of continental thinning in baffin bay margins: insights from svartenhuk halvø peninsula basalts, west greenland. lithos 334–335, 117–140. https://doi.org/10.1016/j.lithos.2019.03.017 andreasen, f. 1981: sedimentological observations in cretaceous and tertiary rocks in the northern part of the west greenland sedimentary basin. rapport grønlands geologiske undersøgelse 105, 26–27. https://doi.org/10.34194/rapggu.v105.7746 bonow, j.m. 2005: re-exposed landforms in the disko region, west greenland – disregarded data for estimation of glacial erosion and uplift modelling. geomorphology 72, 106–127. https://doi. org/10.1016/j.geomorph.2005.05.006 bonow, j.m., japsen, p., lidmar-bergström, k., chalmers, j.a. & pedersen, a.k. 2006: cenozoic uplift of nuussuaq and disko, west greenland – elevated erosion surfaces as uplift markers of a passive margin. geomorphology 80, 325–337. https://doi.org/10.1016/j.geomorph.2006.03.006 chalmers, j.a. & pulvertaft, t.c.r. 2001: development of the continental margins of the labrador sea: a review. in: wilson, r.c.l. et al. 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53.9592 395173 7934433 907 71° 29.6013' n 053° 57.270' w 71.49336 53.9545 395368 7934981 base 401 71° 29.7219' n 053° 55.887' w 71.49536 53.9315 396194 7935165 7 peak 1309 m top flow 1275 71° 44.2424' n 054° 20.678' w 71.73737 54.3446 383070 7962881 flow 694 71° 44.1487' n 054° 18.739' w 71.73581 54.3123 384190 7962645 flow 569 71° 43.7773' n 054° 18.362' w 71.72962 54.3060 384371 7961943 hyaloclastite 372 71° 42.7672' n 054° 18.814' w 71.71279 54.3136 384005 7960083 hyaloclastite 296 71° 42.5761' n 054° 18.197' w 71.70960 54.3033 384345 7959708 hyaloclastite 236 71° 42.2035' n 054° 17.984' w 71.70339 54.2997 384432 7959009 hyaloclastite 227 71° 42.2230' n 054° 17.575' w 71.70372 54.2929 384672 7959033 hyaloclastite 217 71° 42.2535' n 054° 17.052' w 71.70422 54.2842 384980 7959073 hyaloclastite 208 71° 42.2414' n 054° 16.670' w 71.70402 54.2778 385202 7959038 hyaloclastite 193 71° 42.3096' n 054° 15.982' w 71.70516 54.2664 385610 7959143 hyaloclastite 188 71° 42.2356' n 054° 15.530' w 71.70393 54.2588 385866 7958991 hyaloclastite 181 71° 42.2771' n 054° 15.252' w 71.70462 54.2542 386032 7959059 base on sediment 169 71° 42.4850' n 054° 14.581' w 71.70808 54.2430 386445 7959424 9 peak 1078 m top 1089 71° 39.6959' n 054° 22.319' w 71.66160 54.3720 381643 7954494 978 71° 39.2889' n 054° 20.536' w 71.65481 54.3423 382643 7953681 929 71° 38.8403' n 054° 19.561' w 71.64734 54.3260 383167 7952817 873 71° 38.7421' n 054° 18.583' w 71.64570 54.3097 383729 7952603 454 71° 38.9028' n 054° 16.998' w 71.64838 54.2833 384673 7952850 345 71° 38.6420' n 054° 14.541' w 71.64403 54.2424 386084 7952288 302 71° 38.3777' n 054° 13.408' w 71.63963 54.2235 386721 7951762 base 51 71° 38.3167' n 054° 11.974' w 71.63861 54.1996 387554 7951604 10 taseraarsuit base 385 71° 46.8924' n 054° 26.917' w 71.78154 54.4486 379720 7968005 455 71° 47.0278' n 054° 27.267' w 71.78380 54.4544 379532 7968268 566 71° 47.1832' n 054° 27.122' w 71.78639 54.4520 379632 7968552 top nerutusoq mb 723 71° 47.3462' n 054° 26.589' w 71.78910 54.4432 379959 7968836 750 71° 47.3790' n 054° 26.539' w 71.78965 54.4423 379991 7968896 1037 71° 47.6918' n 054° 26.172' w 71.79486 54.4362 380237 7969464 top 1164 71° 47.8158' n 054° 25.409' w 71.79693 54.4235 380693 7969669 12 point 470 m top 403 71° 27.2084' n 053° 58.434' w 71.45347 53.9739 394462 7930571 289 71° 26.9113' n 053° 58.599' w 71.44852 53.9766 394338 7930024 208 71° 26.7809' n 053° 58.358' w 71.44635 53.9726 394468 7929775 base 29 71° 26.6616' n 053° 57.903' w 71.44436 53.9651 394726 7929540 13 saviit south top 579 71° 24.8203' n 054° 20.276' w 71.41367 54.3379 381309 7926813 509 71° 24.8696' n 054° 18.899' w 71.41449 54.3150 382129 7926859 217 71° 24.3754' n 054° 17.630' w 71.40626 54.2938 382831 7925901 base 33 71° 24.4074' n 054° 15.346' w 71.40679 54.2558 384187 7925887 14 saviit north top 601 71° 26.3664' n 054° 20.801' w 71.43944 54.3467 381157 7929700 345 71° 26.1009' n 054° 19.105' w 71.43501 54.3184 382133 7929152 base 99 71° 26.0300' n 054° 16.887' w 71.43383 54.2814 383438 7928948 22 tasiusaq top 804 71° 31.2318' n 054° 47.320' w 71.52053 54.7887 366047 7939655 622 71° 31.0606' n 054° 45.854' w 71.51768 54.7642 366889 7939284 base 137 71° 31.0651' n 054° 44.449' w 71.51775 54.7408 367718 7939240 24 anilaarfissuaq top 841 71° 30.8326' n 054° 18.234' w 71.51388 54.3039 383129 7937908 base 244 71° 30.8580' n 054° 15.736' w 71.51430 54.2623 384603 7937875 29 qiterlikassak top 1041 71° 39.9905' n 054° 43.758' w 71.66651 54.7293 369148 7955779 1012 71° 39.4229' n 054° 42.084' w 71.65705 54.7014 370062 7954665 base 438 71° 38.5019' n 054° 39.780' w 71.64170 54.6630 371304 7952874 appendix 1 http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 119 of 121 appendix 1 svartenhuk profile coordinates altitude geographical coordinates, wgs 84 utm zone 22* no profile name note m latitude longitude lat. °n long. °w easting northing 4 kakilisaat top 936 71° 29.3011' n 053° 57.553' w 71.48837 53.9592 395173 7934433 907 71° 29.6013' n 053° 57.270' w 71.49336 53.9545 395368 7934981 base 401 71° 29.7219' n 053° 55.887' w 71.49536 53.9315 396194 7935165 7 peak 1309 m top flow 1275 71° 44.2424' n 054° 20.678' w 71.73737 54.3446 383070 7962881 flow 694 71° 44.1487' n 054° 18.739' w 71.73581 54.3123 384190 7962645 flow 569 71° 43.7773' n 054° 18.362' w 71.72962 54.3060 384371 7961943 hyaloclastite 372 71° 42.7672' n 054° 18.814' w 71.71279 54.3136 384005 7960083 hyaloclastite 296 71° 42.5761' n 054° 18.197' w 71.70960 54.3033 384345 7959708 hyaloclastite 236 71° 42.2035' n 054° 17.984' w 71.70339 54.2997 384432 7959009 hyaloclastite 227 71° 42.2230' n 054° 17.575' w 71.70372 54.2929 384672 7959033 hyaloclastite 217 71° 42.2535' n 054° 17.052' w 71.70422 54.2842 384980 7959073 hyaloclastite 208 71° 42.2414' n 054° 16.670' w 71.70402 54.2778 385202 7959038 hyaloclastite 193 71° 42.3096' n 054° 15.982' w 71.70516 54.2664 385610 7959143 hyaloclastite 188 71° 42.2356' n 054° 15.530' w 71.70393 54.2588 385866 7958991 hyaloclastite 181 71° 42.2771' n 054° 15.252' w 71.70462 54.2542 386032 7959059 base on sediment 169 71° 42.4850' n 054° 14.581' w 71.70808 54.2430 386445 7959424 9 peak 1078 m top 1089 71° 39.6959' n 054° 22.319' w 71.66160 54.3720 381643 7954494 978 71° 39.2889' n 054° 20.536' w 71.65481 54.3423 382643 7953681 929 71° 38.8403' n 054° 19.561' w 71.64734 54.3260 383167 7952817 873 71° 38.7421' n 054° 18.583' w 71.64570 54.3097 383729 7952603 454 71° 38.9028' n 054° 16.998' w 71.64838 54.2833 384673 7952850 345 71° 38.6420' n 054° 14.541' w 71.64403 54.2424 386084 7952288 302 71° 38.3777' n 054° 13.408' w 71.63963 54.2235 386721 7951762 base 51 71° 38.3167' n 054° 11.974' w 71.63861 54.1996 387554 7951604 10 taseraarsuit base 385 71° 46.8924' n 054° 26.917' w 71.78154 54.4486 379720 7968005 455 71° 47.0278' n 054° 27.267' w 71.78380 54.4544 379532 7968268 566 71° 47.1832' n 054° 27.122' w 71.78639 54.4520 379632 7968552 top nerutusoq mb 723 71° 47.3462' n 054° 26.589' w 71.78910 54.4432 379959 7968836 750 71° 47.3790' n 054° 26.539' w 71.78965 54.4423 379991 7968896 1037 71° 47.6918' n 054° 26.172' w 71.79486 54.4362 380237 7969464 top 1164 71° 47.8158' n 054° 25.409' w 71.79693 54.4235 380693 7969669 12 point 470 m top 403 71° 27.2084' n 053° 58.434' w 71.45347 53.9739 394462 7930571 289 71° 26.9113' n 053° 58.599' w 71.44852 53.9766 394338 7930024 208 71° 26.7809' n 053° 58.358' w 71.44635 53.9726 394468 7929775 base 29 71° 26.6616' n 053° 57.903' w 71.44436 53.9651 394726 7929540 13 saviit south top 579 71° 24.8203' n 054° 20.276' w 71.41367 54.3379 381309 7926813 509 71° 24.8696' n 054° 18.899' w 71.41449 54.3150 382129 7926859 217 71° 24.3754' n 054° 17.630' w 71.40626 54.2938 382831 7925901 base 33 71° 24.4074' n 054° 15.346' w 71.40679 54.2558 384187 7925887 14 saviit north top 601 71° 26.3664' n 054° 20.801' w 71.43944 54.3467 381157 7929700 345 71° 26.1009' n 054° 19.105' w 71.43501 54.3184 382133 7929152 base 99 71° 26.0300' n 054° 16.887' w 71.43383 54.2814 383438 7928948 22 tasiusaq top 804 71° 31.2318' n 054° 47.320' w 71.52053 54.7887 366047 7939655 622 71° 31.0606' n 054° 45.854' w 71.51768 54.7642 366889 7939284 base 137 71° 31.0651' n 054° 44.449' w 71.51775 54.7408 367718 7939240 24 anilaarfissuaq top 841 71° 30.8326' n 054° 18.234' w 71.51388 54.3039 383129 7937908 base 244 71° 30.8580' n 054° 15.736' w 71.51430 54.2623 384603 7937875 29 qiterlikassak top 1041 71° 39.9905' n 054° 43.758' w 71.66651 54.7293 369148 7955779 1012 71° 39.4229' n 054° 42.084' w 71.65705 54.7014 370062 7954665 base 438 71° 38.5019' n 054° 39.780' w 71.64170 54.6630 371304 7952874 appendix 1 (continued) svartenhuk profile coordinates geographical coordinates, wgs 84 utm zone 22* no profile name note altitude m latitude longitude lat. °n long. °w easting northing 30 ivissukkat top 1000 m 955 71° 39.5928' n 054° 35.675' w 71.65988 54.5946 373826 7954754 itinnerat 951 71° 39.2945' n 054° 34.525' w 71.65491 54.5754 374466 7954160 825 71° 38.7064' n 054° 33.465' w 71.64511 54.5577 375021 7953032 base 430 71° 38.5045' n 054° 32.340' w 71.64174 54.5390 375657 7952619 31 aputituut top (trig. point) 1195 71° 34.1699' n 054° 49.769' w 71.56950 54.8295 364952 7945198 qaqqaat 976 71° 33.9223' n 054° 49.013' w 71.56537 54.8169 365366 7944711 893 71° 33.7554' n 054° 47.497' w 71.56259 54.7916 366237 7944345 587 71° 33.6988' n 054° 46.228' w 71.56165 54.7705 366975 7944193 base 360 71° 33.8835' n 054° 45.503' w 71.56472 54.7584 367423 7944509 36b arfertuarsuk base 52 71° 30.5771' n 055° 13.955' w 71.50962 55.2326 350288 7939483 top of trachyte 86 71° 30.6259' n 055° 14.711' w 71.51043 55.2452 349849 7939605 top of trachyte 167 71° 30.9653' n 055° 14.968' w 71.51609 55.2495 349742 7940245 gully 172 71° 30.9041' n 055° 15.044' w 71.51507 55.2507 349689 7940135 gully 133 71° 31.2278' n 055° 17.025' w 71.52046 55.2837 348567 7940818 topographic top 110 71° 31.1758' n 055° 17.616' w 71.51960 55.2936 348212 7940746 stratigraphic top 52 71° 31.1226' n 055° 18.828' w 71.51871 55.3138 347492 7940698 37 nuup qaava top 867 72° 05.1707' n 055° 17.568' w 72.08618 55.2928 352736 8003797 base 31 72° 04.6531' n 055° 23.878' w 72.07755 55.3980 349064 8003098 38 qaqqap qaa top 1154 72° 09.1463' n 054° 37.346' w 72.15244 54.6224 376150 8009664 1011 72° 08.5581' n 054° 38.422' w 72.14263 54.6404 375472 8008609 base 279 72° 07.5396' n 054° 40.446' w 72.12566 54.6741 374204 8006789 40 pannertuup top 1385 72° 17.8153' n 053° 40.214' w 72.29692 53.6702 409398 8024064 qaqqaa base 739 72° 17.6842' n 053° 42.957' w 72.29474 53.7160 407837 8023890 41b skalø south top, western gully 463 71° 49.9775' n 055° 32.700' w 71.83296 55.5450 341970 7976260 plateau edge, w. gully 371 71° 49.8377' n 055° 32.820' w 71.83063 55.5470 341881 7976006 base, western gully 0 71° 49.6648' n 055° 32.382' w 71.82775 55.5397 342110 7975666 profile shift eastern gully 51 71° 49.8523' n 055° 31.807' w 71.83087 55.5301 342469 7975989 base, eastern gully 0 71° 49.8340' n 055° 31.725' w 71.83057 55.5287 342514 7975951 42 marraarnaq top 479 72° 18.7330' n 055° 15.003' w 72.31222 55.2500 355981 8028850 base 37 72° 19.7348' n 055° 15.173' w 72.32891 55.2529 356016 8030715 44c kuugaartorfik highest 1185 71° 58.5361' n 054° 31.628' w 71.97560 54.5271 378251 7989776 799 71° 58.4615' n 054° 33.040' w 71.97436 54.5507 377431 7989685 433 71° 58.6189' n 054° 34.455' w 71.97698 54.5742 376635 7990025 lowest 62 71° 58.5665' n 054° 36.955' w 71.97611 54.6159 375192 7990014 top sediments 454 71° 58.6322' n 054° 34.361' w 71.97720 54.5727 376691 7990047 base sediments 317 71° 58.5882' n 054° 35.091' w 71.97647 54.5849 376266 7989990 47 peak 1430 m highest 1329 71° 59.5814' n 054° 19.829' w 71.99302 54.3305 385139 7991330 lowest 777 71° 59.0298' n 054° 21.343' w 71.98383 54.3557 384213 7990355 47a peak 1625 m highest 1469 71° 57.6395' n 054° 02.906' w 71.96066 54.0484 394677 7987209 lowest 1223 71° 57.4732' n 054° 03.392' w 71.95789 54.0565 394381 7986914 48b nerutusoq lowest, in river 108 71° 55.1077' n 054° 30.278' w 71.91846 54.5046 378656 7983367 hyaloclastite in river 293 71° 53.5627' n 054° 28.632' w 71.89271 54.4772 379439 7980444 hyaloclastite in river 334 71° 53.0149' n 054° 29.236' w 71.88358 54.4873 379031 7979447 hyaloclastite in river 352 71° 52.9096' n 054° 29.621' w 71.88183 54.4937 378798 7979265 hyaloclastite in river 485 71° 50.4418' n 054° 31.237' w 71.84070 54.5206 377596 7974738 hyaloclastite in river 504 71° 49.7162' n 054° 32.427' w 71.82860 54.5405 376829 7973432 hyaloclastite in river 502 71° 49.2440' n 054° 33.446' w 71.82073 54.5574 376186 7972590 highest, on mount. side 658 71° 49.5137' n 054° 35.658' w 71.82523 54.5943 374935 7973167 50 umiiarfik south top 873 71° 52.4722' n 054° 58.579' w 71.87454 54.9763 362021 7979491 150 71° 53.2864' n 054° 58.868' w 71.88811 54.9811 361954 7981012 base 121 71° 53.3771' n 054° 57.310' w 71.88962 54.9552 362864 7981121 51 tunuarsuk top 1278 71° 47.5299' n 054° 49.737' w 71.79216 54.8289 366546 7969988 1157 71° 47.3509' n 054° 50.254' w 71.78918 54.8376 366224 7969675 base 456 71° 46.7082' n 054° 48.302' w 71.77847 54.8050 367281 7968411 http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 120 of 121 appendix 1 (continued) svartenhuk profile coordinates geographical coordinates, wgs 84 utm zone 22* no profile name note altitude m latitude longitude lat. °n long. °w easting northing 52 nuuit qaqqaat top sub-profile 1 905 71° 50.8945' n 055° 13.008' w 71.84824 55.2168 353483 7977131 east side of mountain 725 71° 50.5362' n 055° 13.771' w 71.84227 55.2295 352995 7976498 base sub-profile 1 534 71° 50.2978' n 055° 13.415' w 71.83830 55.2236 353170 7976041 top sub-profile 2 423 71° 50.0397' n 055° 12.185' w 71.83400 55.2031 353848 7975512 268 71° 49.7904' n 055° 11.832' w 71.82984 55.1972 354020 7975036 base sub-profile 2 244 71° 49.7304' n 055° 11.858' w 71.82884 55.1976 353998 7974925 55 aputituut top 848 71° 35.5437' n 054° 52.583' w 71.59240 54.8764 363464 7947853 801 71° 36.0127' n 054° 52.180' w 71.60021 54.8697 363755 7948708 663 71° 36.3566' n 054° 51.903' w 71.60594 54.8650 363959 7949336 base 487 71° 36.9067' n 054° 52.098' w 71.61511 54.8683 363910 7950364 58 amitsoq base 43 71° 48.6321' n 055° 24.281' w 71.81054 55.4047 346658 7973403 167 71° 48.2264' n 055° 23.601' w 71.80377 55.3934 346997 7972621 top sub-profile 1 412 71° 47.7829' n 055° 25.403' w 71.79638 55.4234 345892 7971875 base sub-profile 2 367 71° 47.4048' n 055° 22.625' w 71.79008 55.3771 347452 7971056 top 610 71° 47.0412' n 055° 22.043' w 71.78402 55.3674 347741 7970357 59 qooruusaq strat. lowest 253 71° 43.5862' n 055° 22.416' w 71.72644 55.3736 347059 7963965 132 71° 43.0601' n 055° 25.459' w 71.71767 55.4243 345216 7963119 138 71° 43.0894' n 055° 26.076' w 71.71816 55.4346 344861 7963200 79 71° 43.3524' n 055° 28.348' w 71.72254 55.4725 343575 7963785 strat. highest 124 71° 43.5403' n 055° 29.373' w 71.72567 55.4896 343004 7964178 60 qooruusaq strat. lowest 71 71° 43.0127' n 055° 27.895' w 71.71688 55.4649 343792 7963136 80 71° 42.7501' n 055° 28.665' w 71.71250 55.4778 343307 7962682 107 71° 42.3854' n 055° 28.073' w 71.70642 55.4679 343601 7961980 120 71° 42.2702' n 055° 28.118' w 71.70450 55.4686 343560 7961769 strat. highest 176 71° 41.3399' n 055° 26.374' w 71.68900 55.4396 344449 7959968 61 qooruusaq strat. lowest 34 71° 44.0873' n 055° 31.890' w 71.73479 55.5315 341616 7965302 58 71° 44.0032' n 055° 32.022' w 71.73339 55.5337 341528 7965152 strat. highest 73 71° 43.8406' n 055° 31.482' w 71.73068 55.5247 341819 7964827 62 sigguk 1. base, coast in nw 0 71° 40.9634' n 055° 51.152' w 71.68272 55.8525 329949 7960385 2. coast se of 1. 0 71° 40.7685' n 055° 49.812' w 71.67948 55.8302 330702 7959961 3. coast se of 2. 73 71° 40.8071' n 055° 49.385' w 71.68012 55.8231 330957 7960012 4. coast se of 3. 0 71° 40.6040' n 055° 47.336' w 71.67673 55.7889 332121 7959540 5. base vertical part 32 71° 40.3537' n 055° 44.234' w 71.67256 55.7372 333894 7958933 6. top 640 71° 40.9337' n 055° 44.507' w 71.68223 55.7418 333819 7960021 64a paannivik top of 64a 1292 71° 57.9423' n 054° 58.282' w 71.96570 54.9714 362862 7989629 base of 64a 611 71° 57.5581' n 055° 03.915' w 71.95930 55.0652 359577 7989132 base of 64b 0 71° 55.8438' n 055° 05.582' w 71.93073 55.0930 358402 7986017 197 71° 56.1749' n 055° 05.099' w 71.93625 55.0850 358722 7986612 423 71° 56.3526' n 055° 04.461' w 71.93921 55.0743 359111 7986917 top of 64b 703 71° 56.5799' n 055° 02.653' w 71.94300 55.0442 360180 7987268 66 saattut 'base' (coastal profile) 0 72° 13.1650' n 055° 56.399' w 72.21942 55.9400 331796 8020317 (fladøerne) 'top' 28 72° 18.4205' n 055° 55.364' w 72.30701 55.9227 333181 8030010 easternmost heli-reconnaissance nunatak 566505 1739 72° 04.2366' n 051° 54.493' w 72.07061 51.9082 468803 7997087 65 salliaruseq sample no. and notes 566511, highest 1144 72° 25.0812' n 053° 47.968' w 72.41802 53.7995 405661 8037755 566512 1093 72° 25.1172' n 053° 48.676' w 72.41862 53.8113 405262 8037885 566513 1080 72° 25.0080' n 053° 49.033' w 72.41680 53.8172 405049 8037672 566514 1089 72° 24.9396' n 053° 49.172' w 72.41566 53.8195 404943 8037565 566515 72° 24.8256' n 053° 49.528' w 72.41376 53.8255 404764 8037350 566516 993 72° 24.7962' n 053° 49.564' w 72.41327 53.8261 404725 8037241 566517 981 72° 24.7836' n 053° 49.535' w 72.41306 53.8256 404725 8037241 566518 964 72° 24.7590' n 053° 49.788' w 72.41265 53.8298 404590 8037247 566519 918 72° 24.5148' n 053° 50.131' w 72.40858 53.8355 404367 8036811 http://www.geusbulletin.org www.geusbul let in.org larsen & larsen 2022: geus bulletin 50. 8295. https://doi.org/10.34194/geusb.v50.8295 121 of 121 appendix 1 (continued) svartenhuk profile coordinates geographical coordinates, wgs 84 utm zone 22* no profile name note altitude m latitude longitude lat. °n long. °w easting northing 566520 918 72° 24.5148' n 053° 50.131' w 72.40858 53.8355 404367 8036811 566521 895 72° 24.5058' n 053° 49.783' w 72.40843 53.8297 404564 8036690 566522 826 72° 24.3120' n 053° 49.826' w 72.40520 53.8304 404548 8036355 572001 710 72° 24.2460' n 053° 50.627' w 72.40410 53.8438 404071 8036266 572002 710 72° 24.2460' n 053° 50.627' w 72.40410 53.8438 404071 8036266 572003 700 72° 24.2364' n 053° 50.660' w 72.40394 53.8443 404071 8036266 572004 700 72° 24.2364' n 053° 50.660' w 72.40394 53.8443 404071 8036266 572005 680 72° 24.2508' n 053° 50.872' w 72.40418 53.8479 403936 8036273 572006 670 72° 24.2664' n 053° 50.942' w 72.40444 53.8490 403903 8036274 572007 620 72° 24.2382' n 053° 51.251' w 72.40397 53.8542 403734 8036282 572008. lowest 590 72° 24.2292' n 053° 51.458' w 72.40382 53.8576 403599 8036289 67 qooruusaq sample no. and notes 2019 568604 above trachyte 179 71° 39.1572' n 055° 21.750' w 71.65262 55.3625 346837 7955767 568601 trachyte 191 71° 39.1206' n 055° 21.668' w 71.65201 55.3611 346899 7955650 568602 below trachyte 121 71° 39.0090' n 055° 21.647' w 71.65015 55.3609 346883 7955428 568603 below previous 119 71° 38.9952' n 055° 21.600' w 71.64992 55.3600 346918 7955425 * utm zone 22 coordinates as used for the other onshore parts of the nuussuaq basin. same flow http://www.geusbulletin.org lithostratigraphy, geology and geochemistry of the tertiary volcanic rocks on svartenhuk halvø and abstract introduction previous investigations of the svartenhuk halvø area geological setting general geology structures basin area northern extension of the cretaceous boundary fault system northern and eastern areas methods field work, sample profiles and map compilation geochemistry nomenclature vaigat formation kakilisaat member nerutusoq member nunavik member svartenhuk formation kuugaartorfik member tunuarsuk member nuuit member skalø member naqerloq formation naqerloq formation on svartenhuk halvø arfertuarsuk member intrusions geochemistry and petrology of the volcanic rocks identification of crustal contamination and geochemical enrichment vaigat formation nunavik member nerutusoq member kakilisaat member svartenhuk formation tunuarsuk member nuuit member skalø member discussion naqerloq formation arfertuarsuk intrusions primary magmas, melting conditions and development with time vaigat formation svartenhuk formation naqerloq formation enriched mantle components sporadic addition of enriched lithospheric material an enriched asthenospheric mantle component melting and mixing relations in the svartenhuk formation concluding remarks acknowledgements additional information references appendix 1 figures fig. 1 simplified geological map showing the regional and tectonic setting of the nuussuaq basin in fig. 2 simplified geological map of the nuussuaq basin. light colours: sea-covered areas. cbf: creta fig. 3 stratigraphic scheme and 40ar-39ar radiometric ages (in ma) for the volcanic rocks in the nuu fig. 4 index map of geological maps covering svartenhuk halvø and the surrounding areas. maps with n fig. 5 geological sketch map of west greenland between 71° and 73°n. the northern and eastern extent fig. 6 geological map of svartenhuk halvø and the northern volcanic areas. geology south fig. 7 place names used in the text. names in red are also profile names; their numbers are shown in fig. 8 locations and numbers of all studied profiles through the volcanic succession. profiles with fig. 9 digital elevation model (from the greenland ice mapping project) of svartenhuk halvø with loc fig. 10 profiles through the volcanic succession in southern svartenhuk halvø. vertical scale is m a fig. 11 profiles through the volcanic succession in central and western svartenhuk halvø. vertical s fig. 12 profiles through the volcanic succession in northern svartenhuk halvø. vertical scale is m a fig. 13 profiles through the volcanic succession north of svartenhuk halvø. vertical scale is m a.s. fig. 14 the coastal section at narsinganersua (kap cranstown). location in fig. 8. the faulted lava fig. 15 section along the north coast of tasiusap imaa, 15-20 km north-east of kap cranstown and eas fig. 16 section along the south coast of svartenhuk halvø from illerusat qaqqaat to akunnerit. locat fig. 17 section along the south coast of svartenhuk halvø from saviit to ulissat. location in fig. 8 fig. 18 section along the east coast of southern svartenhuk halvø from maniiseqqut in the south to n fig. 19 the south coast of svartenhuk halvø between saviit and maniiseqqut (location in fig. 7). the fig. 20 faulted lava flows of the kakilisaat (kak) and nunavik (nun) members of the vaigat formation fig. 21 the three members of the vaigat formation in the east wall of the mountain kakilisaat qaqqaa fig. 22 the three members of the vaigat formation on the western slope of the mountain kakilisaat qa fig. 23 pillow lava with close-lying pillows up to 1 m in diameter developed in the lower part of th fig. 24 elongate pillows reminiscent of seals on the shore. kakilisaat member, schade øer, south of fig. 25 pillow breccia with abundant, up to metre-sized pillows, pillow fragments and lava fragments fig. 26 map showing flow directions as indicated by dips of foreset-bedding in hyaloclastites of the fig. 27 the vaigat formation west of the siuteqqut kuuat valley, looking west. in the foreground sub fig. 28 the middle part of the type profile (48b, fig. 9) for the nerutusoq member. the location is fig. 29 brown hyaloclastites (dark) of the nerutusoq member interdigitated with grey hyaloclastites fig. 30 picrites typical of the nunavik member forming a thick succession of grey, crumbling, compou fig. 31 the northernmost exposures of the vaigat formation on the innerit peninsula. the formation h fig. 32 the nunavik member close to reference profile 7. the succession of thin, grey, subaerial pic fig. 33 well-exposed lavas of the nunavik member in the large flexure zone that runs sse from usuit fig. 34 subaerial lava flows of the upper part of the nunavik member of the vaigat formation (nun) o fig. 35 conglomerate with large clasts of volcanic rocks in a matrix of reddish volcanogenic sandsto fig. 36 overlapping flow lobes in a compound pahoehoe lava flow. note the numerous zeolite-filled ve fig. 37 lava succession comprising alternating groups of thick, massive sheet flows and compound pah fig. 38 typical coarse hyaloclastite with unsorted pillow fragments in a matrix of glass grains and fig. 39 the upper part of the nunavik member (nun) containing a crimson-red bed c. 8 m thick of weld fig. 40 nunavik member in central svartenhuk halvø around the taseraarsuit lake (blue oval), viewed fig. 41 the most complete exposed section through the svartenhuk formation, resting on the vaigat fo fig. 42 tio₂ (wt%) for profiles through the svartenhuk and naqerloq formations. the prof fig. 43 type section for the kuugaartorfik member in profile 44c, kuugaartorfik. vf: vaigat formatio fig. 44 yellow and dark grey sediments of the kuugaartorfik member dipping west, well exposed in the fig. 45 siltstone-dominated sediments of the kuugaartorfik member on the north coast of the innerit fig. 46 white quartzofeldspathic, poorly consolidated sandstones of the kuugaartorfik member overlyi fig. 47 lava flows of the tunuarsuk member comprising intervals with massive sheet flows intersperse fig. 48 the lava succession in the steep east wall of qaqqap qaa, eastern innerit peninsula. profile fig. 49 volcanic rocks of the svartenhuk formation overlying precambrian metasediments of the karrat fig. 50 irregular, columnar-jointed lava sheets invading volcanogenic sediments (vs). lower part of fig. 51 lava flow with a basal colonnade of coarse columns along the shore, changing upwards to slen fig. 52 the boundary between lava flows of the tunuarsuk member (tun) and the more brownish lava flo fig. 53 the nuuit and skalø members of the svartenhuk formation at nuuit qaqqaat, looking nnw. the w fig. 54 characteristic brown, massive lava flows of the nuuit member overlying thinner and more vari fig. 55 thick lava flow with a 20-25 m high colonnade with bent column tops. nuuit member at asungas fig. 56 multi-tiered columns in a thick, ponded lava flow, probably the same flow as in fig. 55. hei fig. 57 three successive lava flows (la) of the nuuit member banked up against a steep hill of preca fig. 58 volcanic plug feeding one of the lowest lava flows in the nuuit member near the boundary to fig. 59 the upper lava flows of the nuuit member and the conformable boundary to the skalø member at fig. 60 light grey lava flows of the skalø member overlying massive brown flows of the nuuit member fig. 61 lava flows of the skalø member with characteristic yellow-tinted, light top zones, overlying fig. 62 heterogeneous volcaniclastic sediments that form the uppermost horizon of the skalø member. fig. 63 the northern outlier of the arfertuarsuk trachyte in the uppermost reaches of the qooruusaq fig. 64 lava flows of the paleocene nuuit and skalø members of the svartenhuk formation conformably fig. 65 the arfertuarsuk trachyte flow exposed on the beach, east side of the northern end of the ar fig. 66 feldspar-rich xenolith, probably cognate, of syenite in the arfertuarsuk trachyte flow. leng fig. 67 fine-grained, dark grey inclusion in trachyte from the outlier in qooruusaq, cut into two. t fig. 68 fine-grained, dark grey vein cutting the arfertuarsuk trachyte flow. east side of the northe fig. 69 distribution and directions of dykes on svartenhuk halvø as identified on aerial photographs fig. 70 an intrusion of magnesian basalt into lava flows of the tunuarsuk member. a: the entire expo fig. 71 th/nb vs. nb/zr diagram showing discrimination between crustally contaminated and geochemica fig. 72 major-element variation diagrams for rocks of the vaigat formation. note the mgo-rich charac fig. 73 incompatible trace-element variation diagrams for rocks of the vaigat formation. the nunavik fig. 74 transition-element variation diagrams for rocks of the vaigat formation. note the apparent l fig. 75 ree and multi-element diagrams for representative rocks of the vaigat formation. different c fig. 76 major-element variation diagrams for rocks of the svartenhuk formation. crustally contaminat fig. 77 incompatible trace-element variation diagrams for rocks of the svartenhuk and naqerloq forma fig. 78 transition-element variation diagrams for rocks of the svartenhuk and naqerloq formations an fig. 79 ree and multi-element diagrams for representative rocks of the svartenhuk formation. the dif fig. 80 major-element variation diagrams for basalts of the naqerloq formation. a group of low-ti ba fig. 81 ree and multi-element diagrams for representative rocks of the naqerloq formation. the diffe fig. 82 ree and multi-element diagrams for dykes. the distinction between dykes of the svartenhuk an fig. 83 ree cross plots and mantle melting models for the primary melts of the volcanic rocks of the fig. 84 nb/y vs. zr/y diagrams. a: vaigat formation. black contour line is the field of the ordlinga fig. 85 ce/y vs. zr/nb diagram demonstrating the compositional differences between the svartenhuk fo fig. 86 multi-element diagram for selected samples showing variable patterns of element enrichment. fig. 87 stratigraphic variation up-section of lan/ndn (chondrite normalised) in eight densely sample tables table 1 tunuarsuk member thicknesses and number of flows with greater and lesser than 8 wt% mgo table 2 nuuit member thicknesses, number of lava flows and average flow thicknesses table 3 skalø member thicknesses, number of lava flows and average flow thicknesses table 4 ⁴⁰ar/³⁹ar age determinations for the arfertuarsuk trachyte flow table 5 chemical analyses of rocks from the vaigat formation table 6 chemical analyses of rocks from the svartenhuk formation table 7 notes on analysed samples from the naqerloq formation and the delta-1 drill hole table 8 chemical analyses of dykes, sills and samples dredged from baffin bay geological survey of denmark and greenland bulletin 13, 2007, 37-40 37 recently discovered volcaniclastic rocks of andesitic composition form major parts of the mid-archaean, amphibolite facies supracrustal belts at qussuk, on bjørneøen and on part of storø in western godthåbsfjord (fig. 1). these rocks are interpreted as an island arc that represents the onset of the magmatic accretion of the akia terrane 3070 ma ago; this terrane is the north-westernmost of several archaean tec tono-stratigraphic terranes in the nuuk region, which were all amalgamated by 2720 ma (cf. hollis et al. 2006). the presence of the arc in the akia terrane points to similarities between high-grade orthogneiss-amphibolite associations in west greenland and lower-grade granite-greenstone terrains of other archaean cratons e.g. in canada and western australia. volcaniclastic rocks belonging to the ancient arc have been subject to intense synvolcanic, hydrothermal alteration associated with gold-copper mineralisation especially in parts of the qussuk area. another important gold prospect occurs on central storø, which is currently being explored by nunaminerals a/s (knudsen et al. 2007 – this volume). this contribution presents new field observations from some of the best preserved parts of the ancient arc at qussuk and on bjørneøen, while it remains unclear if the volcano-sedimentary associations and their gold mineralisation at qussuk, bjørneøen and the nearby storø share a common midarchaean geological history. the age and setting of the ancient arc the eastern akia terrane comprises c. 3060–3000 ma tonalitic to trondhjemitic orthogneisses, isoclinally folded panels derived from the older andesitic arc, and granites mobilised from the orthogneisses during a late-kinematic thermal event at c. 2980 ma which also led to granulite facies p-t conditions west of qussuk and bjørneøen (garde 1997, 2007; garde et al. 2000). andesitic metavolcanic and volcano-sedimentary rocks were first reported from the eastern akia terrane by garde (1997) and smith (1998). however, volcanic textures are mostly very poorly preserved, and it was not appreciated then, how widespread the andesitic arc is, and that it forms large parts of the supracrustal belts exposed at qussuk and on bjørneøen. recent age determinations of volcanic zircon grains from central bjørneøen show that the age of the andesitic arc is 3071 ± 1 ma (sample 479827, garde 2007). the arc is thus marginally older than its orthogneiss host, in agreement with recently observed intrusive contacts into the supracrustal belt (fig. 2). similar field relationships in the adjacent qussuk area further document that the orthogneiss precursors intruded into the arc. however, zircon from volcaniclastic rocks here, on strike with those on bjørneøen, yields metamorphic ages of 2990–2970 ma which coincide with the thermal maximum and associated fluid movement; only a few zircon cores approach the true volcanic age of c. 3070 ma pre-metamorphic hydrothermal alteration with gold in a mid-archaean island arc, godthåbsfjord, west greenland adam a. garde, henrik stendal and bo møller stensgaard © geus, 2007. geological survey of denmark and greenland bulletin 13, 37–40. available at: www.geus.dk/publications/bull fig. 1. simplified map of north-western godthåbsfjord with supracrustal belts (green), gold anomalies, and locations of figs 2–8. modified from hollis (2005). (garde 2007). zircon grains from yet another volcano-sedimentary rock on central bjørneøen (fig. 3a) yielded a range of 2908–2742 ma ages with a cluster around 2825 ma (sample 479745, hollis 2005 p. 55). the 2825 ma cluster was interpreted as the depositional age, and a complex tectonic model for central bjørneøen was proposed with thrust-stacking of supracrustal rocks and orthogneisses of different ages and origins at the eastern margin of the akia terrane. however, further field observations in 2006 uncovered that this sample locality lies within the same volcaniclastic sequence as the sample dated at 3071 ma. with the established intrusive contacts of orthogneisses dated at c. 3065– 3050 ma (fig. 2; hollis 2005 pp. 30–39), this tectonic model is no longer tenable, as the supracrustal belt is older than the orthogneisses and its contacts are not tectonic. furthermore, a duplicate sample collected at the locality of fig. 3a by the first author in 2006 yielded a metamorphic upb zircon age of 2986 ± 3.6 ma (13 stubby, very indistinctly zoned zircon grains with th/u <0.01; unpublished ion probe data, may 2007). hence, the zircon grains allegedly extracted from sample 479745 seem to have been incorrectly labelled during the sample preparation. volcaniclastic and volcano-sedimentary rocks the relict volcano-sedimentary arc has been intensely de formed and metamorphosed at middle to upper amphibolite grade and now mostly consists of monotonous, fine-grained, schistose andesitic rocks besides intrusive bodies of mafic–ultramafic rocks. however, small pockets of low-strain rocks with distinct volcaniclastic textures are locally well preserved on central bjørneøen and in the qussuk area, particularly in the cores of fold hinges. andesitic rocks containing volcanic clasts with fiamme textures have been identified on central bjørneøen and north-eastern qussuk (fig. 3), as well as rare, more or less undeformed volcano-sedimentary depos its a few tens of metres across with well-preserved sedimentary structures (fig. 4). the recognition of the volcaniclastic environments is important, because they document widespread explosive volcanism and hence shallow subaqueous or subaerial volcanic activity in an island arc. unfortunately, it has not been possible to reconstruct individual volcanic edifices due to the general poor state of preservation. synvolcanic hydrothermal alteration hydrothermally altered aluminous and siliceous rocks, commonly associated with disseminated iron sulphides and gold mineralisation, occur in several parts of the relict arc (hollis 2005; garde 2007). unequivocal evidence that the hydro thermal alteration took place in unconsolidated volcaniclastic rocks prior to deformation and metamorphism was found 38 fig. 2. intrusive contacts at the footwall (a) and hanging wall (b) of the supracrustal belt on central bjørneøen (see fig. 1 for locations). fig. 3. volcaniclastic rocks on central bjørneøen (a) and at north-eastern qussuk (b) with fiamme textures. pen 8 mm thick. for locations see fig. 1. in 2006 within the hinge zone of a large fold on the northeastern coast of qussuk (fig. 1). the lower part of fig. 5 displays several nearly undeformed beds of fineto coarsegrained tuff (or tuffite) with granitic-pegmatitic partial melt veins. the thickest bed displays right way-up graded bedding. the contact to the overlying bed is rusty weathering and contains sporadic iron sulphides and small garnets (<5 mm), and has clearly been affected by hydrothermal activity. the very localised alteration furthermore indicates that the percolating fluid used the unconsolidated bedding contact as a convenient passageway. there are many other examples of rocks with more widespread and more intense alteration. most of these are also intensely deformed and difficult to recognise as hydrothermally altered, but exceptions do occur. figure 6 shows smooth, grey metatuff hosting an irregular, interfingering, vein-like system of coarser, rusty weathering and crumbling metatuff. this is hydrothermally altered and now contains abundant quartz, biotite, garnet and iron sulphide. rusty weathering and variably gold-mineralised rocks, altered by synvolcanic, hydrothermal fluids and now rich in quartz, biotite, garnet and commonly sillimanite, also occur east and north of qussuk in tracts up to several kilometres long, which have previously been mapped as aluminous metasedimentary rocks. figure 7 shows a football-sized enclave of unaltered and undeformed amphibolitic metatuff, which is preserved inside a lens several metres across of garnet-rich rock (hydrothermally altered and metamorphosed tuff ). the most intensely altered andesitic rocks now consist of massive sillimanite-fuchsite quartzite, from which most major and trace elements have been leached out except si, al and ti, and very immobile elements such as ga, zr and rareearth elements. such rocks form the top of ivisaat mountain north of qussuk (fig. 8), where they form a layer c. 5–10 m thick in the core of an isoclinal fold, surrounded by garnetrich and sulphidic, less altered andesitic metavolcanic rocks with widespread gold mineralisation in the ppb range (the authors’ unpublished data). this very specific type of hydro fig. 4. graded volcano-sedimentary rock with early brittle/ductile deformation. central bjørneøen (see fig. 1 for location). fig. 5. bedded metatuff at north-eastern qussuk (see fig. 1 for location) showing a primary contact between mafic and intermediate beds (red arrow) and a narrow rusty hydrothermal alteration zone along primary bedding contact (blue arrow). fig. 6. interfingering system of early hydrothermal alteration in andesitic (meta)tuff at north-eastern qussuk (see fig. 1 for location). glacial striations have produced a faint oblique ‘foliation’ visible in the lower left of the exposure. fig. 7. enclave of unaltered and undeformed metavolcanic amphibolite (arrow), surrounded by a lens of hydrothermally altered, now garnetrich rock of similar origin. peninsula east of qussuk (see fig. 1 for location). hammershaft is c. 85 cm long. 39 thermal alteration points to acid leaching under low pressure, a process which is very characteristic of epithermal, high-level hydrothermal systems in modern andesitic arcs, which are commonly associated with gold and copper mineralisation (e.g. sillitoe & hedenquist 2003). a low-pressure hydrothermal system such as found today near the tops of andesitic arcs is a prerequisite for wholesale leaching by strong acids, be cause higher pressures prevent the dissociation of their hydrogen ions (sillitoe & hedenquist 2003). conclusions the c. 3070 ma andesitic arc in the eastern akia terrane and its slightly younger intrusive counterparts of tonalitic ortho gneisses, combined with previous structural evidence of early crustal shortening in most of the akia terrane, point to the existence of a convergent plate-tectonic system in the north atlantic craton, where subduction of oceanic crust and partial melting of the subducted slab occurred at least 3070 ma ago. the identification of the arc complex substantiates previous ideas that the orthogneisses in the akia terrane are products of slab melting in a convergent plate-tectonic setting. in a wider context this also implies that the typical archaean high-grade orthogneiss-amphibolite associations in west greenland may not represent plate-tectonic environments distinct from the granite-greenstone associations found in most other archaean cratons, but simply expose deeper sections of the same convergent systems. the relict arc hosts widespread hydrothermally altered rocks and associated gold (-copper) mineralisation. newly discovered field relationships show that the hydrothermal systems predated deformation and metamorphism. further more, the peculiar mineralogical composition of the altered rocks combined with their gold mineralisation suggest that the hydrothermal alteration was synvolcanic and epithermal, and characteristic of the arc itself. references garde, a.a. 1997: accretion and evolution of an archaean high-grade grey gneiss–amphibolite complex: the fiskefjord area, southern west greenland. geology of greenland survey bulletin 177, 114 pp. garde, a.a. 2007: a mid-archaean island arc complex in the eastern akia terrane, godthåbsfjord, southern west greenland. journal of the geological society (london) 164, 565–579. garde, a.a., friend, c.r.l., nutman, a.p. & marker, m. 2000: rapid matu ration and stabilisation of middle archaean continental crust: the akia terrane, southern west greenland. bulletin of the geological society of denmark 47, 1–27. hollis, j.a. (ed.) 2005: greenstone belts in the central godthåbsfjord region, southern west greenland: geochemistry, geochronology and petrography arising from 2004 field work, and digital map data. danmarks og grønlands geologiske undersøgelse rapport 2005/42, 215 pp. hollis, j.a., frei, d., van gool, j.a.m., garde, a.a. & persson, m. 2006: using zircon geochronology to resolve the archaean geology of southern west greenland. geological survey of denmark and greenland bulletin 10, 49–52. knudsen, c., van gool, j.a.m., østergaard, c., hollis, j.a., rinkjørgensen, m., persson, m & szilas, k. 2007: gold hosting supracrustal rocks on storø, west greenland: petrogenesis and structural setting. geological survey of denmark and greenland bulletin 13, 17–20. sillitoe, r.h. & hedenquist, j.w. 2003: linkages between volcanotectonic settings, ore-fluid compositions, and epithermal precious metal deposits. in: simmons, s.f. & graham, i. (eds): volcanic, geothermal, and ore-forming fluids; rulers and witnesses of processes within the earth. society of economic geologists special publication 10, 315–343. smith, g.m. 1998: geology and mineral potential of the bjørneøen supracrustal belt, nuukfjord, west greenland. unpublished report, nunaoil a/s, 13 pp. (in archives of geological survey of denmark and greenland, geus report file 21649). 40 author’s address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: aag@geus.dk fig. 8. folded sillimanite-fuchsite quartzite with disseminated iron sulphides, an intensely hydrothermally altered and metamorphosed rock of andesitic origin. ivisaat mountain north of qussuk (see fig. 1 for location). hammershaft is c. 45 cm long. prins et al. 2025: geus bulletin 52. 8371. https://doi.org/10.34194/sj67gw16 1 of 11 review article a review of subsurface geosystems and de-risking offshore construction in the danish north sea lasse tésik prins*1 , katrine juul andresen1,3 , matthew owen2 , paul c. knutz1 1department of land and marine geology, geological survey of denmark and greenland (geus), aarhus, denmark; 2global maritime, london, uk; 3department of geoscience, aarhus university, aarhus, denmark abstract the renewable energy transition has increased the demand for offshore construction in the danish north sea energy sector. this development underpins the need for further investigation of potential geological hazards and associated risks to avoid accidents involving people, the environment or infrastructure. a scientific approach to de-risking requires an understanding of the seabed and the buried geosystems. understanding geosystems is the first step in the de-risking process of offshore construction. in this study, we review three key geosystem elements in the danish north sea, represented by (1) shallow stratigraphy and geomorphology, (2) glacial tectonics and salt movement and (3) subsurface fluid migration. we summarise the current state of knowledge of these geosystem elements and identify multiple risks associated with each geosystem in the region. such investigations are critical for understanding the geotechnical behaviour of the subsurface and identifying and de-risking of potential geohazards during the construction of future energy developments in the danish north sea region. *correspondence: lkp@geus.dk received: 19 feb 2024 revised: 29 nov 2024 accepted: 18 feb 2025 published: 14 apr 2025 keywords: buried tunnel valleys, engineering geology, glacial tectonics, marine geology, late quaternary abbreviations: cpt: cone penetration test dns: danish north sea eez: exclusive economic zone geus: geological survey of denmark and greenland lgm: last glacial maximum pgm: preliminary ground model geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: margaret dolan (geological survey of norway), kaskela anu (geological survey of finland) reviewed by: daniel hepp (nihk, germany), mark coughlan (icrag, ireland), nicola dakin (british geological survey, uk) funding: see page 8 competing interests: see page 9 additional files: see page 9 1. introduction the danish north sea (dns) is an important asset that provides a suite of important functions and services to danish and international societies. these include not only fisheries and cargo transportation but also important geosystem services such as oil and gas production, co2 storage and offshore renewable energy. in addition to these services, the dns contains important habitats for a variety of birds, fish and marine mammals (danish maritime authority 2023). the renewable energy transition will impose increased offshore construction pressure, including windfarms, cable routes and energy islands to the dns. this places a demand on marine spatial planning, not just for the dns but the entire north sea region, which includes the german, british, dutch, belgian and norwegian exclusive economic zones (eezs; cotterill et al. 2017a; fleischer et al. 2022; petrie et al. 2022; danish maritime authority 2023). the risks associated with these large-scale offshore construction projects are potentially immense, in terms of both geotechnical issues, health and safety and environmental risks, including construction failure, underwater noise, suspended sediments, pollution and changes to marine habitats (le et  al. 2014; degraer et al. 2020; mooney et al. 2020). adding to the complexity of the energy transition is the uncertainty concerning the rentability of offshore wind projects. sound economic models are going to be crucial for reaching a 6 gw output from offshore wind energy to meet a danish reduction target of 70% compared to 1991 (møllgaard et al. 2024). https://doi.org/10.34194/sj67gw16 https://orcid.org/0000-0002-2824-8345 https://orcid.org/0000-0001-8029-3234 https://orcid.org/0000-0001-7119-4656 https://orcid.org/0000-0001-5188-4254 mailto:lkp@geus.dk https://creativecommons.org/licenses/by/4.0/deed.ast prins et al. 2025: geus bulletin 52. 8371. https://doi.org/10.34194/sj67gw16 2 of 11 geusbulletin.org to facilitate marine planning of the dns in the energy transition era that builds on safe and sustainable approaches, a comprehensive understanding of the underlying marine geosystems is required. due to hydrocarbon prospecting and extraction over the last five decades (adegbamigbe et al. 2022), there is already significant knowledge on the deep geology. a similar knowledge base for the shallow geosystems, i.e. the upper 500 m below the seafloor, is only just starting to emerge. a geosystem is closely related to the term geosystem services. this has been defined as either ‘the direct result of the planet’s geodiversity’ or as providing ‘benefits specifically resulting from the subsurface’ (frisk et al. 2022). in terms of de-risking offshore construction, we have adopted a more specific definition. thus, geosystems in this context constitute geological features that influence offshore construction. through reviewing published literature, this paper describes how knowledge of near-surface geology in the offshore environment of the dns can help to de-risk the offshore construction process. we focus on three key geosystem elements represented by shallow stratigraphy and geomorphology, glacial tectonics and salt migration and subsurface fluid migration (table 1). 1.1. geological background the north sea is an epicontinental sea, bordered by the uk to the west, norway and denmark to the east and germany, the netherlands and belgium to the south (fig. 1). the north sea basin was initiated as a rift system in the early triassic, which terminated in the paleocene (ziegler 1992). up to 3000 m of cenozoic sediments have accumulated in the basin ( cameron et  al. 1987; huuse & clausen 2001; gołedowski et  al. 2012; ottesen et  al. 2014). the thickness of the sediments representing the quaternary period (last 2.6  million years) may be up to 800 m in the central north sea (nielsen et  al. 2008; ottesen et  al. 2014; phillips et  al. 2017). these sediments are typically heterogenous and have been deposited, reworked and deformed largely as a result of glacial processes associated with the cyclic expansion and decay of large ice sheets in north-west europe (knudsen & sejrup 1993; hughes et al. 2016; rea et al. 2018; batchelor et al. 2019; kirkham et  al. 2022). during the elsterian (c. 500–400 kyr bp) and saalian (c. 380–130 kyr bp) glacial periods, the entire dns was covered by ice (van der vegt et al. 2012), which locally resulted in erosion into older palaeogene or cretaceous deposits. prominent remnants of these glaciations are expressed as buried tunnel valleys (huuse & lykke-andersen 2000; benvenuti et al. 2018; prins et al. 2020) or as glaciotectonic complexes ( andersen et al. 2005; winsemann et al. 2020). conversely, the more recent weichselian glaciation (c. 117–11.5 kyr bp) covered only the northern and western parts of the dns (fig. 2d; hughes et al. 2016). the last glacial maximum (lgm; 22–18 kyr bp) represents a phase of major expansion of the fennoscandian ice sheet inducing large variations in geomorphology and sediment distribution across the region. areas close to the weichselian ice margin are often marked by an increase in geological complexity, which may involve interlayering of till, meltwater sediments and proglacial lake infill deposits (fig. 2c). these features may be preserved within buried valleys and other paleo-landscape depressions such as those formed by glaciotectonic deformation (moreau & huuse 2014; prins & andresen 2019). during and after the lgm, widespread marsh deposits formed in a boreal semi-submerged landscape of the german north sea (coughlan et al. 2018). a similar landscape evolution has also been suggested for the southwestern part of the dns based on shallow seismic and acoustic data (prins & andresen 2019; andresen et  al. 2022) although paleo-environmental constraints on landscape development around the last low-stand are sparse. following the last deglaciation, from about 11,000 years bp, most of the dns was presumably above sea level, forming a low relief landscape with lakes and bogs commonly infilling topographic depressions (coughlan et  al. 2018). as the post-glacial landscape became inundated by rising sea-levels, multiple channels were formed in connection with riverine drainage systems (e.g. elbe paleo-valley), which gradually transformed into estuaries (hepp et  al. 2017, 2019; prins & andresen 2019; andresen et al. 2022). continuation of the holocene transgression meant that by about table 1 summary of geosystems and their associated risks and methods of identification. geosystem associated risks risk-reducing investigations shallow stratigraphy and geomorphology unpredicted lateral variations in soil behaviour. improved regional stratigraphic models. glacial tectonics and salt movement unpredicted lithology, potential fluid migration, variation in geotechnical properties. mapping past ice movements and their influence on the sediments; describing overburden above salt structures. fluid migration potential blowouts, changes in soil cohesion/ strength. understanding shallow fluid migration paths and mechanisms; mapping shallow gas. https://doi.org/10.34194/sj67gw16 http://www.geusbulletin.org/ prins et al. 2025: geus bulletin 52. 8371. https://doi.org/10.34194/sj67gw16 3 of 11 geusbulletin.org 9.3 kyr bp, the dns was subject to full marine conditions and influenced by strong tidal currents that initiated the deposition of large sand banks, presently known as little fisher bank and the jylland bank (leth 1996; fig. 3). these features, like many other shallow areas of the dns, remain hydrographically dynamic with mobile sand units forming the modern seabed (anthony & leth 2002; nørgaard-pedersen & rödel 2021). our knowledge of north sea geosystems that identifies the foundation zone for offshore constructions is rooted in the complex quaternary strata, involving multiple phases of deposition, erosion and glacial loading and unloading. although several studies on the quaternary stratigraphy in the north sea have emerged in recent years (le bot et  al. 2005; rijsdijk et  al. 2005; cotterill et al. 2017b; coughlan et al. 2018; prins & andresen 2019; petrie et al. 2022), the quaternary succession of the dns remains poorly constrained, both in terms of chronostratigraphy, lithological variation and geotechnical properties. closing this knowledge-gap requires densely spaced and high-resolution data coverage aimed at mapping the shallow geosystems and would provide information that can de-risk offshore services now and in the future. 1.2. de-risking in a geosystem context risk, in the context of a geosystem, can be regarded as an assessment of the probability and consequence resulting in a negative impact produced by a specific geosystem element (copping et al. 2020). this could be the likelihood and consequence of a ship grounding in areas with varying water depths, or the likelihood and consequence of a punch-through failure for a jack-up rig, installed on an undiscovered buried valley. a buried tunnel valley poses no risk to a ship passing over it, but if it results in a subsurface failure below a jack-up rig, the consequences can potentially be fatal (bienen et al. 2015). consequences may also be financial, for example, as recently experienced in the neart na gaoithe offshore windfarm in the uk sector, where the likelihood of wind turbine generator foundation installation failing due to rockhead variability had not been foreseen and had a financial consequence of hundreds of millions of euros (watts et al. 2021). such examples illustrate that 6100000 55° 2’ 12’’ 55°56’ 6’’ 56° 49’ 59’’ 4° 5’ 32’’ 5° 43’ 28’’ 7° 21’ 39’’ –50 m –50 m –50 m –50 m –50 m 4° 5’ 32’’ 5° 43’ 28’’ 7° 21’ 39’’ 55 °5 6’ 6 ’’ energy island renewable energy development o�shore renewable energy development o�shore installations mean elevation (m) –10 –55 –30 no eez uk eez dk eez de eez nl eez fig. 1 map overview of the danish north sea (dns) and some of the offshore activities in the area (danish maritime authority 2023). base map projection: etrs89 utm zone 32n. background map: emodnet bathymetry. eez: exclusive economic zone. no: norway. dk: denmark. de: germany. nl: netherlands. uk: united kingdom. red lines indicate eezs. https://doi.org/10.34194/sj67gw16 http://www.geusbulletin.org/ prins et al. 2025: geus bulletin 52. 8371. https://doi.org/10.34194/sj67gw16 4 of 11 geusbulletin.org risk is closely related to the interaction between humans and geosystems. understanding geosystems is the first step in the de-risking process of offshore construction. this typically involves a desktop study, which describes the potential risks at a specific location, based on existing data and knowledge from the literature (e.g. owen et al. 2020). desktop studies highlight areas that require further investigation and help to inform potential risks associated with the area of interest. during the initial phases of site investigation, the first step is to generate a representative preliminary ground model (pgm) for the investigation area, which is typically based on any existing geophysical, geological and geotechnical data. the pgms will then subsequently be developed further to form a fully integrated ground model based on additional site investigations and the collection of site-specific data (cook et al. 2014). the pgm will also aid in the survey design to optimise ship time and thus reduce costs. further advantages from having a high level of 2000 m 100 m 200 m 50 m north sea n orw egian channel db lb bergen norway 55ºn 0ºn s n s multiple water remobilised sediments fluvial sediments deposits of baltic ice lake shallow channel in�ll outwash phase ii outwash phase i lacustrine deposits subglacial till and pre-quaternary deposits 1 km 10 ms peak trough 1 km 10 ms 40 80 120 (m s) 40 80 120 (m s) 53 52 bh/51 50 49 48 47 46 1 2 3 4 1 2 3 4 ? onlap & downlapthrustsdecollement surface 100 200 300 400 tw t ( m s) 100 200 300 400 tw t ( m s) sw sw ne ne hill hole 500 m 500 m seismic re�ector (0.78 ma) brunhes/matuyama seabed multiple gas front ’chimney’ stratigraphy visible beneath pockmark due to venting of gas or �uid 2 m deep depression visible on bathymetry and innomar data bathymetry data innomar dataa b c d fig. 2 overview of different categories of geosystems reported in previous publications. a: subsurface fluid migration: flow features represented on innomar sub-bottom profiler data and multibeam echosounder data as pockmarks and trapped gas (modified from owen et al. 2021). b: glacial tectonics and salt movement: shown by a seismic profile and model representation of a glaciotectonic complex (modified from bendixen et al. 2017). twt: two-way travel time. c: shallow stratigraphy and geomorphology: here visualised on various geophysical data from the baltic sea (modified from bellwald et al. 2023). d: extent of a proposed glacial lake (orange polygon) during the last glacial maximum (lgm), which has potentially deposited problematic soft clays over a large area (green shaded areas indicate db (dogger bank) and lb (ling bank); modified from hjelstuen et al. 2018). https://doi.org/10.34194/sj67gw16 http://www.geusbulletin.org/ prins et al. 2025: geus bulletin 52. 8371. https://doi.org/10.34194/sj67gw16 5 of 11 geusbulletin.org understanding of the geosystems include reducing the risk of project delays due to unforeseen ground conditions. these are extremely costly and not only delay the project itself but also reduce the chance of meeting political goals such as reducing carbon emissions. 2. geosystem elements and their potential risks this paper focuses on the potential risks associated with the following three geosystem elements in the dns (fig. 2): 1. shallow stratigraphy and geomorphology (fig. 2c and 2d). 2. glacial tectonics and salt movement (fig. 2b) 3. subsurface fluid migration (fig. 2a) 2.1. shallow stratigraphy and geomorphology an understanding of the shallow stratigraphy (typically down to 200–500 m below seabed) is the basic framework for gaining information on the subsurface geotechnical properties. knowing the lateral and vertical variation in lithology and mapping subsurface stratigraphic boundaries are essential for foundation design for large constructions such as wind turbine generators or artificial islands (fig. 2c). in the inner danish waters, a recent desktop study identified a thick succession of weakly consolidated glaciomarine clays (jensen & bennike 2022). this resulted in the windfarm project south of hesselø to be paused and demanded new site investigations to assess the subsurface geological constraints in the area. in the dns, numerous site surveys have been carried out in relation to offshore construction, but efforts to synthesise the results between sites have been sparse (prins & andresen 2021; petrie et al. 2024). integrating geotechnical data with robust stratigraphic models has the potential to reduce both risks and costs in relation to offshore construction (velenturf et al. 2021; petrie et al. 2024) and help avoid delaying large-scale offshore construction projects. ice-dammed lakes typically form massive deposits of soft clay, which can pose a hazard for the development of offshore windfarms. a phase of ice-dammed lake development during the last glacial maximum has been suggested for the dns area (hjelstuen et al. 2018; 55° 2’ 12’’ 55°56’ 6’’ 56° 49’ 59’’ 4° 5’ 32’’ 5° 43’ 28’’ 7° 21’ 39’’ 4° 5’ 32’’ 5° 43’ 28’’ 7° 21’ 39’’ 55 °5 6’ 6 ’’ elbe palaeo-valley outline tunnel valleys (2d) dogger bank outline lgm ice extent proposed glacial lake extent buried valleys (3d) eez borders salt structures seabed sediments mud and sandy mud muddy sand sand gravel and coarse sand till/diamicton quaternary clay & silt jyske rev lille fiskerbanke fig. 3 overview of the different geosystems at various stratigraphic depths in the danish north sea (dns). eez: exclusive economic zone. lgm: last glacial maximum. https://doi.org/10.34194/sj67gw16 http://www.geusbulletin.org/ prins et al. 2025: geus bulletin 52. 8371. https://doi.org/10.34194/sj67gw16 6 of 11 geusbulletin.org fig.  2d). the geographical extent of this lake phase is bound by uncertainties although recent observations support a wider presence across the central and southern dns (andresen et al. 2022; knutz et al. 2022). further work is needed to constrain these clay-rich units in terms of lateral extent, thickness and geotechnical properties. understanding the subsurface geology requires knowledge of the stratigraphic units, their distribution and their lithological variation. however, in order to generate a predictive subsurface model, the geological history and depositional environment need to be constrained, which require the geomorphology and sedimentological processes of the subsurface to be interpreted (cotterill et al. 2017a). the shallow subsurface in the north sea contains multiple different glacial geomorphologies, including prominent positive landforms such as eskers and moraine ridges (dove et  al. 2017; emery et  al. 2019; mellett et  al. 2020) and paleo-coastline deposits, e.g. aggradational bars, which are well known in onshore denmark. buried negative landforms that are less expressive in the terrestrial terrain are commonly observed in the offshore seismic profiles as various forms of channels and troughs. these predominantly erosional features range from small troughs formed locally in a tidal paleo marsh setting (coughlan et  al. 2018) to kilometre-scale buried valleys formed by fluvial or subglacial processes. the elbe paleo-valley forms a major depressional feature of composite erosional channels that intersects the dns in a sse–nnw direction (lonergan et al. 2006; stewart et al. 2013; moreau & huuse 2014; ottesen et al. 2014; cotterill et al. 2017b; prins & andresen 2019; emery et al. 2020; prins et al. 2020). depending on their origin and subsequent geological evolution, e.g. the character of sedimentary infill, buried valleys pose different risks for offshore construction, particularly deployment of jack-up rigs and design of appropriate wind turbine foundations. peats have frequently been reported from within buried valleys as well as the surrounding fluvial plains (coughlan et al. 2018; hepp et al. 2019), and these pose a risk for cable routings, as they increase the risk of overheating (bellwald et al. 2024). by assembling and integrating all the available sub-surface information, geological models can be produced that describe spatial variations in buried geomorphology, litho-stratigraphy and depositional environments. subsequently, geotechnical information can be added, or inferred, to generate a ground model that will be used for spatial site planning of offshore installations. an example is the dogger bank windfarm area, where intensive data collection and the integration of geological and geotechnical data have been suggested to potentially reduce the need for drilling, thus lowering the cost of the site survey investigations (cotterill et al. 2017a). 2.2. glacial tectonics and salt movement the glaciation history of the dns has led to variable glaciotectonic effects on the geosystems. deformation in a subglacial or proglacial environment occurs when the weight of a moving ice sheet exerts a lateral stress component in the subsurface strata, causing failure or brittle deformation, which propagates through the ice-contact zone (andersen et al. 2005). this results in thrusting and folding of the pre-existing strata, which may lead to stacking and repetition of the sedimentary sequences (bennet & glasser 2009) introducing geological heterogeneity and unpredictability in the area (fig. 2b). glaciotectonic complexes are found throughout the north sea providing evidence of ice-marginal processes during the last and previous glaciations (andersen et al. 2005; larsen & andersen 2005; bendixen et  al. 2017; cotterill et al. 2017b; pedersen & boldreel 2017; owen et  al. 2020). some of the most well-studied glaciotectonic deformation structures are found on the island of mors in denmark, where diatomite and ash layers show extensive folding of paleocene–eocene deposits (klint & pedersen 1995). glaciotectonic thrust complexes alter the existing stratigraphy through deformation with potentially discontinuous sedimentary sequences as a result, such as allochthonous slabs of fine-grained material in a sandy matrix. this introduces increased heterogeneity and facies unpredictability in the area. a concrete example of glaciotectonically induced heterogeneity can be found in the jammerbugt area, where the upper cretaceous chalk units have been deformed by glaciotectonic activity, leaving a depression that was subsequently filled with eemian weichselian deposits (pedersen & boldreel 2017). similar effects are seen in the british north sea sector, where the rugged surface of thrust complexes has facilitated deposition of fine-grained material within lakes or ponds, some of which may contain organic-rich deposits, e.g. peat or gyttja (cotterill et al. 2017a). for the final risk assessment, variability in lithology and structural character induced by glacial tectonics need to be integrated into the geotechnical ground model (velenturf et al. 2021). beyond the negative effects of glacial deformation and substratum complexities, ice loading may also be beneficial for offshore foundations as it enhances burial compaction and may lead to over-consolidation expressed by high shear strength and sediment stiffness (le et al. 2014). deformation of the near surface sediments can also occur as a result of salt movement (rank-friend https://doi.org/10.34194/sj67gw16 http://www.geusbulletin.org/ prins et al. 2025: geus bulletin 52. 8371. https://doi.org/10.34194/sj67gw16 7 of 11 geusbulletin.org & elders 2004). and there is evidence of quaternary faulting related to these salt movements (huuse et  al. 2001). the  most prominent influence these deep salt structures pose on offshore construction is through the faulting of the near-surface sediments as well as fluid migration through these faults. 2.3. subsurface fluid migration subsurface fluid migration is a naturally occurring process in sedimentary basins, principally driven by sediment compaction, decomposition of organic matter and development of localised pressure gradients (judd & hovland 2009). evidence of fluid migration within geosystems is seen as crater-like depressions, or pockmarks, on the seafloor (lohrberg et al. 2020; andresen et al. 2021) or in the form of gas chimneys, pipes and buried pockmarks on seismic data (fig. 2a; cartwright et al. 2007; andresen et al. 2008; andresen 2012; moss et al. 2012). seaor lake-floor pockmarks on bathymetry data (reusch et al. 2015; lohrberg et al. 2020; andresen et al. 2021) or pockmarks at the present day land surface (bogoyavlensky et al. 2020) are documented from many sites globally – particularly sites located in hydrocarbon-prone sedimentary basins such as the dns, where enhanced fluid flow, commonly associated with salt-induced geological structures, is prevalent (huuse et  al. 2010; knutz 2010) or along permafrost or gas hydrate regions (walter anthony et al. 2012). seafloor pockmarks may pose a risk to offshore installations. if fluid expulsion is active, overpressurised pore fluids may yield low sediment stability or fluidisation of sediments below the crater or rim of the pockmark (hovland et al. 2002; chuvilin et al. 2020). in the norwegian north sea sector, seafloor pockmarks are a common feature in offshore windfarm development areas, which needs to be assessed in terms of fluid migration activity and potential risks (petrie et al. 2022). locations of focused fluid seepage may also present issues due to marine habitat protection of bubble reefs. these bioherms are formed by chemosynthetic organisms that use methane as an energy source whilst precipitating authigenic carbonate (noble-james et al. 2020). fluid escape and the presence of shallow gas may also occur without a prominent seafloor expression. gas in the shallow subsurface is found across most of the dns, where fine-grained sediments with organic content are present, or where geological conditions are amenable to vertical gas migration (etiope 2009; vielstädte et  al. 2015; petersen & smit 2023). in a de-risking context, it is important to understand the spatial distribution and geological context of subsurface gas accumulations as they can lead to gas blowouts or undermine foundations for offshore installations. thus, understanding the gas migration pathway through geosystems, whether related to natural processes or induced by human activities, such as oil and gas production (hornafius et  al. 1999) or wind turbine foundation is crucial for reducing risk elements in offshore construction (coughlan et al. 2021). 2.4. combined risk elements fluid migration within glaciotectonised areas is a case of combined risk elements that can influence the geotechnical properties within a geosystem. fluid or gas migration may occur along thrust planes of the deformation complex, thus acting as a conduit for deeper fluids, which may reach the subsurface stratum and cause foundation conditions to deteriorate (velenturf et  al. 2021). similar fluid migration issues may be relevant near large salt structures (fig. 3). 3. from geosystems to ground models this paper highlights the importance of understanding geosystem elements that carry potential hazards for future energy developments in the dns. improving this knowledge will reduce physical, environmental and financial risks associated with the expected increase in offshore construction activities. this risk reduction is essential for avoiding accidents, project delays or cancellations for large offshore renewable projects that are a key component in meeting decarbonisation and net zero targets (møllgaard et al. 2024). the geosystems risk analyses should be treated as an integrated approach, since interference with one system can pose new unknown risks in other systems. for instance, although the lithological succession may appear static, the strain and loading induced from site survey drilling and construction or piling may alter the physical properties in the subsurface by compaction and fluid expulsion. mitigation of subsurface interactions occurring during the construction phase requires modelling of subsurface behaviour based on knowledge from the geological models and geosystem elements. integrated models that compile all geophysical, geological and geotechnical data are key for defining the interaction between geosystem services and the risk associated with a specific geosystem (prins & andresen 2021; velenturf et  al. 2021; bellwald et  al. 2023). the model presented by prins & andresen (2021) included scattered data points from the central graben area from 11 site surveys including seven cone penetration tests (cpts), which are 1d geotechnical measurements of the strength of the sediments. the data spread highlights the need for a better regional understanding of the subsurface geology. the development of offshore windfarm sites requires large amounts of geophysical, https://doi.org/10.34194/sj67gw16 http://www.geusbulletin.org/ prins et al. 2025: geus bulletin 52. 8371. https://doi.org/10.34194/sj67gw16 8 of 11 geusbulletin.org geotechnical and geological data to ground truth the engineering properties of the subsurface. and as such, regional ground models can be improved, providing a better foundation for the de-risking process in future offshore construction projects. the improvements shown in the integrated ground models that are currently being produced in relation to large-scale wind energy projects like the ten noorden van de waddeneilanden wind farm zone in the netherlands (see offshorewind.rvo.nl) have a huge potential for extrapolation to wider areas. it also shows how a quantitative approach to integrating geological, geophysical and geotechnical data sets provides additional detailed information that also helps in the de-risking process (karkov et al. 2022). in the irish sea, data collected for ireland’s marine resource program were used to map geotechnical and geological constraints for offshore construction (coughlan et  al. 2020; guinan et  al. 2020). a similar approach could be adopted in the dns where offshore geological data made available from different industry and academic sources allow for qualitative interpretations and early-stage geological models to be established. by leveraging a general understanding of the soil properties, implications for foundation type and design, and potential risks, these models provide a valuable service for decision-makers in marine spatial planning and the offshore energy industry (coughlan et  al. 2020; guinan et al. 2020). data density has improved drastically in some areas of the dns, particularly in association with a proposed energy island (knutz et al. 2022), and as the renewable energy transition progresses, a growing amount of geophysical and geotechnical data will become available. much of these data will be available to the public through the marine raw material database ‘marta’ (geus 2024), which means that an increasingly larger proportion of the dns can be described, for example, involving detailed ground models and stratigraphic schemes, which ultimately allows for a more robust risk analyses. 4. conclusions for the dns, three geosystems potentially containing hazards and forming risks to offshore construction have been identified (table 1). these are: 1. shallow stratigraphy and geomorphology: here, unknown variations in soil strength and lithological variations across the dns pose a risk for offshore construction if it is not resolved through a better general understanding or through site survey investigations. this includes mapping of geomorphological features, which can contribute to large local lithological variations within a stratigraphic unit (e.g. erosional valleys or coarse-grained coastal deposits) but which also helps highlight areas that require further investigation. 2. glacial tectonics: including the postdepositional alteration of the subsurface resulting in unexpected lithological variations (e.g. chalk in a quaternary setting) and potential fluid migration paths, resulting in unexpected geotechnical properties. 3. fluid migration: here, the presence of gas in the sediment causes a variety of potential hazards like expulsion to the water column, loss of sediment cohesion and strength and potential blowouts. this review of geosystems in the dns shows that there are multiple risks associated with each geosystem. it highlights the need for a general understanding of these geosystems and provides information on good practice for understanding and mitigating these risks. because the near-surface geology of the dns is so diverse, there is a need for a broad geological model that can be continuously improved upon. this model should inform the desktop studies and subsequent integrated ground models for offshore construction projects. as we have shown, most regional geosystems and their distribution are known; however, we still lack specific knowledge on how the systems are interlinked, which exact geotechnical properties they have as well as the challenges they pose. as the renewable energy transition continues, more data will need to be gathered about these geosystems, and how to handle the risks associated with them. with increased use of oil and gas technology applied on shallow data, the resolution and accuracy will also gradually increase. to this end, some initiatives have already been established. these include a large-scale mapping project funded by the danish energy agency, to screen danish waters for areas that are unfit for windfarm construction and the noarg project, funded by geocenter denmark (see https://www. geocenter.dk/projekter/2023-2/), which aims to map and improve our understanding of buried valleys in the dns. such projects will increase regional knowledge of the geosystems in the dns and how they might pose a risk to future offshore construction. acknowledgements we would like to acknowledge the danish offshore technology centre at dtu denmark for funding the project seep – building a seabed environmental baseline for platform abandonment. this study is a product of the work carried out in that study. we would also like to thank the three reviewers whose comments significantly improved the manuscript. additional information funding statement this study was funded by the danish offshore technology centre, dtu, denmark. https://doi.org/10.34194/sj67gw16 http://www.geusbulletin.org/ http://offshorewind.rvo.nl https://www.geocenter.dk/projekter/2023-2/ https://www.geocenter.dk/projekter/2023-2/ prins et al. 2025: geus bulletin 52. 8371. https://doi.org/10.34194/sj67gw16 9 of 11 geusbulletin.org author contributions ltp: conceptualisation, formal analysis, methodology, visualisation, writing – original draft, writing – review and editing. kja: conceptualisation, writing – original draft, writing – review and editing. mo: conceptualisation, writing – original draft. pk: conceptualisation, writing – review and editing. competing interests the authors declare no competing interests. additional files none provided. references adegbamigbe, t., biswas, a., blaney, j. & jagtap, a. 2022: the north sea – a long and proud history. 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https://www.upstreamonline.com/energy-transition/saipem-woe-as-soil-and-foundations-issues-hit-work-on-north-sea-wind-farm-installation/2-1-971065 https://doi.org/10.1016/j.quascirev.2019.106068 https://doi.org/10.1016/b978-0-444-89912-5.50007-7 https://doi.org/10.1016/b978-0-444-89912-5.50007-7 a review of subsurface geosystems and de-risking offshore construction in the danish north sea 1. introduction 1.1. geological background 1.2. de-risking in a geosystem context 2. geosystem elements and their potential risks 2.1. shallow stratigraphy and geomorphology 2.2. glacial tectonics and salt movement 2.3. subsurface fluid migration 2.4. combined risk elements 3. from geosystems to ground models 4. conclusions acknowledgements additional information funding statement author contributions competing interests additional files references figures fig. 1 map overview of the danish north sea (dns) and some of the offshore activities in the area (danish maritime authority 2023). base map projection: etrs89 utm zone 32n. background map: emodnet bathymetry. eez: exclusive economic zone. no: norway. dk: denmark. de: germany. nl: netherlands. uk: united kingdom. red lines indicate eezs. fig. 2 overview of different categories of geosystems reported in previous publications. a: subsurface fluid migration: flow features represented on innomar sub-bottom profiler data and multibeam echosounder data as pockmarks and trapped gas (modified from owen et al. 2021). b: glacial tectonics and salt movement: shown by a seismic profile and model representation of a glaciotectonic complex (modified from bendixen et al. 2017). twt: two-way travel time. c: shallow stratigraphy and geomorphology: here visualised on various geophysical data from the baltic sea (modified from bellwald et al. 2023). d: extent of a proposed glacial lake (orange polygon) during the last glacial maximum (lgm), which has potentially deposited problematic soft... fig. 3 overview of the different geosystems at various stratigraphic depths in the danish north sea (dns). eez: exclusive economic zone. lgm: last glacial maximum. table table 1 summary of geosystems and their associated risks and methods of identification. kaskela & kotilainen 2024: geus bulletin 52. 8317. https://doi.org/10.34194/geusb.v52.8317 1 of 8 research article | short quantifying seabed geodiversity of the archipelago sea, baltic sea, finland anu m. kaskela* , aarno t. kotilainen geological survey of finland (gtk), espoo, finland abstract this study investigated the geodiversity of the archipelago sea in the northern baltic sea, focusing on geological features and their spatial distribution. by adapting methods used in previous baltic sea studies, we conducted spatial analyses of geological data sets including bedrock type, seabed substrates and seabed structures. bedrock and substrate data were freely available, while seabed structures were modelled from bathymetry data. geodiversity was quantified using a geodiversity index, which considers the variety of physical elements, roughness and area of the unit. the analyses revealed a diverse seabed environment in the archipelago sea with varying geodiversity throughout the study area. significant features contributing to geodiversity included bedrock fracture and fault zones and large end-moraine formations. similar patterns have been observed in terrestrial areas of finland. the analyses also detected relations between archipelago zonation and geodiversity with areas of open sea more homogeneous than the middle and inner archipelago. this study formally recognises the complexity of the seabed in the archipelago sea and highlights the importance of understanding the geological processes shaping the region. the results can inform maritime spatial planning and sustainable resource management. *correspondence: anu.kaskela@gtk.fi received: 01 apr 2022 revised: 25 aug 2023 accepted: 26 oct 2023 published: 27 feb 2024 keywords: archipelago zonation, geomorphology, marine landscape, seafloor mapping, semi-automated spatial analysis abbreviations: btm: benthic terrain modeler emodnet: european marine observation and data network emma: ecologically significant marine underwater area velmu: finnish inventory programme for underwater marine diversity geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: margaret dolan (ngu, norway), verner brandbyge ernstsen (geus, denmark) reviewed by: zyad al-hamdani (geus, denmark) and one anonymous reviewer. funding: see page 7 competing interests: see page 7 additional files: see page 7 1 introduction geodiversity describes the natural range of geological, geomorphological and soil features along with their assemblages, relationships, properties, interpretations and systems (gray 2004). recognition of geodiversity helps emphasise that nature contains both abiotic and biotic components, and that both geodiversity and biodiversity should be acknowledged as part of an integrated approach to environmental management (gray 2008; crofts 2014). geodiversity assessment involves evaluating geological features and systems using either qualitative or quantitative methods. while research in this field has advanced significantly in recent years, few studies have focused on marine geodiversity (see crisp et al. 2021). this study aimed to expand our understanding of marine geodiversity by applying regional-scale geodiversity assessment methods to a local scale in an area known as the finnish archipelago sea (fig. 1; kaskela et al. 2012; kaskela & kotilainen 2017). the objectives were to document geodiversity and identify environments characterised by high geodiversity and of interest to marine resource management. located along the southwestern coast of finland, the archipelago sea represents one of the most extensive archipelagos in the world (fig. 1). this area has been shaped by geological processes over millions of years, resulting in a rich variety of landscapes, landforms and geological features that make it a hotspot for geodiversity in the baltic sea region. precambrian crystalline basement, glacial erosion and deposition, uplift, extensive coastline and postglacial sedimentary processes all contribute to regionally variable topography and a mosaic of sedimentary environments (e.g. häkkinen 1990; virtasalo et al. 2007; kaskela & kotilainen 2017; kaskela & rinne 2018). https://doi.org/10.34194/geusb.v52.8317 https://orcid.org/0009-0004-0501-8629 https://orcid.org/0000-0002-9347-5907 mailto:anu.kaskela@gtk.fi https://creativecommons.org/licenses/by/4.0/ kaskela & kotilainen 2024: geus bulletin 52. 8317. https://doi.org/10.34194/geusb.v52.8317 2 of 8 geusbulletin.org 2 material and methods we analysed the geodiversity of the archipelago sea, adapting regional methods developed for the baltic sea (kaskela et al. 2012; kaskela & kotilainen 2017). spatial analyses were performed using the arcmap 10.6 platform and the euref fin tm35fin coordinate system. for this study, we primarily used 1:100 000 scale data with the smallest cartographic unit being approximately 0.05 km2. based on prior experience, a pixel size correlating to about 0.25 mm on a map is reasonable for spatial data analyses (e.g. kotilainen et al. 2014). therefore, all data sets were analysed in raster grid format with a pixel size of 25 × 25 m. the geodiversity parameters of bedrock type, seabed substrates and seabed structures, along with thematic scales that were parsed among four to five classes were adopted from a previous study (kaskela & kotilainen 2017). the study area covered approximately 3900 km2 (fig. 1). spatial analyses were focused on marine areas only. the numerous islands in the study area were covered by integrating terrestrial and marine data into continuous layers. 2.1 bedrock bedrock data covering both marine and terrestrial areas at 1:100 000 scale were obtained from geological survey of finland (gtk 2014). we identified four general lithological classes of sandstone, limestone, crystalline rock and rapakivi granite (fig. 2a). 2.2 substrate for marine areas, we used seabed substrate data from the european marine observation and data network (emodnet) at 1:100 000 scale (emodnet geology 2021). ss2 ss1 ss3 200000 250000 300000 66 50 00 0 67 00 00 0 study area salpausselkä minor fault zone major fault zone bathymetry (m) 0 –120 0 10 205 km data: emodnet bathymetry consortium 2020; gtk (2014, 2018); emodnet geology 2021. coordinates: euref fin tm35fin. administrative boundaries from european commission, eurostat/gisco (© eurogeographics). storlandet jurmo rosala nötö gullkrona n fig. 1 location and map of the archipelago sea along the south-west coast of finland in the northern baltic sea. the numbered ss lines in green refer to salpausselkä formations, which are large, marginal sedimentary ridges deposited during glacial standstills of the last deglaciation. the numbers refer to the age order of the formations: ss1 is the oldest, ss2 is of intermediate age and ss3 is the youngest. bathymetry data from emodnet bathymetry consortium (2020). marine and terrestrial bedrock data including faults are from gtk (2014). marine seabed substrate data are from emodnet geology (2021). terrestrial substrate data are from gtk (2018). coordinate system is euref fin tm35fin. https://doi.org/10.34194/geusb.v52.8317 http://www.geusbulletin.org/ kaskela & kotilainen 2024: geus bulletin 52. 8317. https://doi.org/10.34194/geusb.v52.8317 3 of 8 geusbulletin.org the data included five substrate categories of mud to muddy sand, sand, coarse substrate, mixed sediment and bedrock and boulders (fig. 2b; see kaskela et al. 2019 for details of substrate data). for terrestrial areas, we used the superficial deposits of finland data at 1:200 000 scale obtained from the geological survey of finland (gtk 2018). original categories were classified to approximately match the emodnet seabed substrates (kaskela et al. 2019). 2.3 seabed structures five seabed structures including crests, basins, valleys and troughs, plains and slopes were modelled from the elevation and bathymetry model provided by the finnish inventory programme for underwater marine diversity (velmu; fig. 2c; rinne et al. 2014; virtanen et al. 2018). for this purpose, the original 20 m pixel size of the bathymetry model was converted to 25 m pixels. for the modelling of the seabed structures, we adapted the approach presented in lundblad et al. (2006), kaskela et al. (2012) and kaskela & kotilainen (2017). we used the arcmap extension benthic terrain modeler 3.0 (btm; walbridge et al. 2018). we adjusted the neighbourhood sizes to a mesoscale based on studies by reijonen et al. (2008), rinne at al. (2014) and kaskela & rinne (2018). we used an inner radius of 1 cell for both broad and fine scales. an outer radius of 20 cells (500 m) was used for fine scale and an outer radius of 200 cells (5000 m) for broad scale. slopes were classified based on demek (1972) and by defining two degrees as a boundary value. 2.4 quantifying the seabed geodiversity geodiversity was quantified according to a published geodiversity index (serrano & ruiz-flaño 2007; hjort & luoto 2010) as follows: gd eg r sin( ) = × (1) where gd is the geodiversity index, eg is the number of different physical elements (variety or richness) of the fig. 2 seabed geodiversity was quantified based on a: rock type (modified from gtk 2014); b: seabed substrates (emodnet geology 2021); and c: seabed structures. d: background data showing specific geologic features (faults and salpausselkä formations; gtk 2014; emodnet bathymetry consortium 2020). crystalline rocks limestones 0 5 10 20 km a. rock type mud – muddy sand sand coarse mixed bedrock b. seabed substrate0 5 10 20 km 0 5 10 20 km c. seabed structure valleys basins crests plains slopes storlandet jurmo nötö gullkrona rosala d. background data salpausselkä formations fault zones study area 0 5 10 20 km n n n n ss1 ss2 ss3 https://doi.org/10.34194/geusb.v52.8317 http://www.geusbulletin.org/ kaskela & kotilainen 2024: geus bulletin 52. 8317. https://doi.org/10.34194/geusb.v52.8317 4 of 8 geusbulletin.org unit, r is the roughness of the unit, and s represents the area covered by the unit. eg was calculated using the arcmap focal statistics tool. the variety of features in the geological data sets were initially separately calculated per unit area and then summed for the resulting total richness layer as: eg = rock type variety + seabed substrate variety + seabed structure variety (2) the geological data sets were considered equally important, and no weighting was applied. r represents the standard deviation of slope and is interpreted as an effective measure of surface roughness (grohmann et al. 2011). the unit size defined by a circle with a radius of 80 cells (2000 m) was scaled from kaskela & kotilainen (2017). 3 results and discussion results support the consensus that the archipelago sea represents a highly diverse and fragmented seabed environment (see kaskela et al. 2012; kaskela & kotilainen 2017). our analyses indicate geodiversity varying at local scales across the archipelago sea with the geodiversity index (no units) ranging from 0.2 to 5.4 (table 1; see also supplementary file s, for histograms of geodiversity index and its parameters). roughness varies between 0.3 and 5.2 degrees. the total richness varies between 6 and 12 with total richness classes 9 to 11 covering approximately 96% of the area. a study of the baltic sea (kaskela & kotilainen 2017) reported total richness varying from 3 to 14 across the entire basin and richness values for the archipelago sea similar to those reported here. the number of seabed substrates varies between 1 and 5, and areas with three substrates represent the most common class (c. 56%; table 1). structure varies between classes 2 and 5 with most areas falling into class 5 (94%). because crystalline bedrock comprises most of the study area (99%), variability of bedrock appears low. the previous regional-scale study found a higher variety of rock types and seabed structures in the archipelago sea area than in other parts of the baltic sea (kaskela & kotilainen 2017). the areas with the highest geodiversity are located near the islands of storlandet, jurmo and rosala, as well as north of the salpausselkä end-moraine formation ss3, deposited during the last deglaciation (fig. 3). relatively high geodiversity index values appear close to fracture zones and regions with steep slopes (figs 1, 3a and 3c) supporting the interpretation that the structural character of the bedrock generates local-scale geodiversity (kaskela & kotilainen 2017). long-term exposure of fractures and faults in the bedrock and glacial erosion may have further increased pre-glacial relief. in the archipelago sea, these fractures often occur as elongate channels that serve as water-exchange routes between the bothnian sea and the baltic proper. areas of high geodiversity are identified in the salpausselkä areas and particularly in the region north of ss3 (fig. 3). the salpausselkä formations ss1–ss3 appear in seabed surface data as geomorphic ridges and consist of sand and coarse substrate deposits that increase the geodiversity of an area. our geodiversity assessment agrees with that for the terrestrial landscape scale reported by tukiainen & hjort (2021), who found that areas near the salpausselkä zones were among the most diverse. these authors noted that areas where bedrock influences topographic variation on land tend to have higher geodiversity. parallels in marine geodiversity likely stem from the geological history of the region. the baltic sea basin has experienced multiple glaciations over the past c. 3 million years characterised by periods of intense erosion and accumulation of glacio-marine sediment. during the last glacial maximum (c. 20 ka), northern europe and the baltic sea basin were completely covered by an extensive ice sheet. the entire basin had become exposed (deglaciated) by about 10 ka (stroeven et al. 2016). both marine and terrestrial areas show extensive glacial sedimentary cover. glacio-isostatic processes also led to below sea-level submergence of most of finland during this time. in addition to bedrock type and historical glacial processes, geodiversity estimates reflect archipelago zonation. the archipelago sea has been classified into zones grading from coastal areas to open-marine conditions (inner, middle and outer archipelago zones; häyrén 1900; jaatinen 1960; granö et al. 1999). the exposed table 1 geodiversity index statistics. richness/variability is the number of different physical elements in an area defined by a circle with a radius 2000 m. roughness represents the standard deviation of slope. geodiversity index total richness bedrock variability substrate variability seabed structure variability roughness (°) mean 2.3 9.6 1 3.6 4.9 2.4 median 2.3 9 1 3 5 2.4 mode 2.6 9 1 3 5 2.3 minimum 0.2 6 1 1 2 0.3 maximum 5.4 12 2 5 5 5.2 https://doi.org/10.34194/geusb.v52.8317 http://www.geusbulletin.org/ kaskela & kotilainen 2024: geus bulletin 52. 8317. https://doi.org/10.34194/geusb.v52.8317 5 of 8 geusbulletin.org fig. 3 distribution of seabed geodiversity in the archipelago sea, northern baltic sea. a: geodiversity index (no units). index values were calculated based on b: roughness, c: total variability as established from d: rock, e: substrate and f: structure variability. geodiversity index and roughness are represented by their minimum–maximum values. g: background data showing specific geologic features (faults and salpausselkä formations; gtk 2014; emodnet bathymetry consortium 2020). a. geodiversity index 0 10 205 km 5.2 0.3 b. roughness 5.4 0.2 c. total variability 6 7 8 9 10 11 12 d. rock variability 1 2 e. substrate variability 1 2 3 4 5 f. structure variability 2 3 4 5 g. background data storlandet jurmo nötö gullkrona rosala salpausselkä formations fault zones study area n ss1ss2 ss3 https://doi.org/10.34194/geusb.v52.8317 http://www.geusbulletin.org/ kaskela & kotilainen 2024: geus bulletin 52. 8317. https://doi.org/10.34194/geusb.v52.8317 6 of 8 geusbulletin.org outer archipelago areas exhibit less geodiversity compared to the other two zones. in areas outside the fracture and fault zones, the seabed environment of the outer areas was characterised by flat surfaces or basins resulting in low roughness values (fig. 2). visual assessment indicates that areas of high geodiversity occur in the middle and inner archipelago zones as well as in areas with high roughness values (typically greater than 2.8). areas with multiple islands such as rosala, storlandet, nötö and gulkrona (locations in fig. 1) exhibit higher geodiversity due to the dynamic nature of sedimentary processes occurring there as deposition and erosion vary over short distances. as a concept, geodiversity has intrinsic value. it also provides abiotic ecosystem services (gray 2011; gray et  al. 2013). nature is composed of both abiotic and biotic components, and together they can inform integrated environmental-management strategies (gray 2008; crofts 2014). for example, geodiversity may serve as a proxy or support for biodiversity mapping by representing habitats (e.g. kaskela et al. 2017). ecologically significant marine underwater areas (emmas) were designated to enable effective marine spatial planning and sustainable use of finnish marine areas (lappalainen et al. 2019). use of both ecological and geological data identifies emmas in the study area exhibiting high geodiversity that are likely, for example, to host endangered species, rare habitat types or diverse macroalgae and blue mussel communities. geodiversity can therefore help to resolve ecologically significant areas or parameters. geodiversity may also correlate with resource availability and historical significance. results from this study show that areas with high geodiversity, such as rosala, are also associated with early human settlements. during the iron age, an important waterway passed through a nearby area (e.g. fig. 4.1 in mägi 2015). prehistoric inhabitants likely used natural harbours and prioritised strategic and aesthetic aspects of the landscape arising from its geological diversity (e.g. bourassa 1992; gray 2004). the present case study used geodiversity parameters and thematic scales from a previous broad-scale study (kaskela & kotilainen 2017). since the study area mainly consists of crystalline bedrock and the area is predominantly characterised by the same structural class variability, the geodiversity index primarily reflects patterns in substrate variability and roughness. future research should assess how well thematic classifications capture necessary geodiversity elements at this scale. geodiversity parameters and classifications currently vary among quantification studies making it difficult to compare geodiversity results. while we used four to five categories for each parameter, tukiainen & hjort (2021) designated 16 rock types and dolan et al. (2022, this volume) as well as dolan & bjarnadóttir (2023, this volume) designated 10 geomorphon classes. in addition, it is worth noting that the substrate categories themselves are not entirely comparable in terms of diversity. for instance, mixed sediment contains more grain sizes than sand. this difference in intrinsic diversity between substrate categories is not captured by the analysis presented here. to account for intrinsic substrate diversity, dolan et al. (2022, this volume) have suggested quantifying entropy between sediment fraction layers as a possible solution. sediment fraction maps may be estimated from categorical grain-size maps based on indicative values per sediment class. where data permit, grainsize fractions may be modelled directly (e.g. misiuk et al. 2019; mitchell et al. 2019). such an approach may also form a step in the production of categorical grain-size maps. however, direct mapping or modelling of grainsize fractions usually requires laboratory analyses, which is beyond the scope of many mapping initiatives due to cost and time constraints. we recommend that the geodiversity research community works towards standardising geological parameters and their thematic classifications for different spatial scales across different physical elements, which represent different components of geodiversity. this will help to improve the comparability of results, especially among studies addressing similar research questions. 4 conclusions this study systematised and quantified physical characteristics of the archipelago sea. we show that the study area represents a highly diverse and fragmented marine environment with varying geodiversity throughout. the open-water marine areas display greater homogeneity, while middle and inner archipelago zones display greater geodiversity. bedrock fracture and fault zones as well as large end-moraine formations represent significant features contributing to geodiversity. results also identified overlap between terrestrial and marine geodiversity reflective of the recent glacial history of the region. future research should focus on establishing a consensus on the geological parameters and their thematic classifications applied at given spatial scales to better integrate finding across different studies. geological features and geodiversity contribute to the aesthetics, historical significance and ecosystem services in the study area. quantifying these parameters can inform best management practices and sustainable use of marine and coastal areas. recognising unique geological features that shape our landscape can help preserve environmental resources for future generations. https://doi.org/10.34194/geusb.v52.8317 http://www.geusbulletin.org/ kaskela & kotilainen 2024: geus bulletin 52. 8317. https://doi.org/10.34194/geusb.v52.8317 7 of 8 geusbulletin.org acknowledgements the study was conducted as part of the maameri research project ‘strengthening the knowledge base for improving the status of coastal waters in the archipelago sea’. research infrastructure facilities were provided by finmari (finnish marine research infrastructure network) and data were provided by velmu (finnish inventory programme for underwater marine diversity) and emodnet (european marine observation and data network, funded by the european climate, infrastructure and environment executive agency (cinea) through contract easme/emff/2020/3.1.11/lot2/si2.853812_emodnet – geology 5.2). the study also supports implementation of the life-ip biodiversea (enhancing the marine and coastal biodiversity of the baltic sea in finland and promoting the sustainable use of marine resources, life20 ipe/fl/000020) project. we are grateful to eija eräkare for her assistance with some of the analyses. reviewers are thanked for their valuable and constructive feedback. the staff at geus bulletin are acknowledged for their patience, professionalism, and assistance with the manuscript. figures showing depth model or its derivatives have been published with permission ar23694 from the finnish defence office. additional information funding statement this study was funded by geological survey of finland, the projects emodnet (easme/emff/2020/3.1.11/lot2/si2.853812_emodnet – geology 5.2) and maameri as well as the finmari network and the velmu programme. the study also supports implementation of the project life-ip biodiversea (life20 ipe/fl/000020). the project has received funding from the life programme of the european union. the material reflects the views by the authors, and the european commission or the cinea is not responsible for any use that may be made of the information it contains. author contributions amk: modelling, statistical analyses and preparation of the first draft. atk: drafting and revision of the manuscript. competing interests none declared. additional files one supplementary file containing histograms of geodiversity index and its parameters is available at https://doi.org/10.22008/fk2/yi15yl. references bourassa, s.c. 1992: the aesthetics of landscape. london: bellhaven press. crisp, j.r., ellison, j.c. & fischer, a. 2021: current trends and future directions in quantitative geodiversity assessment. progress in physical geography: earth and environment 45(4), 514–540. https://doi. org/10.1177/0309133320967219 crofts, r. 2014: promoting geodiversity: learning lessons from biodiversity. proceedings of the geologists’ association 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(eds.): maritime societies of the viking and medieval world. leeds: maney publishing, 41–61. reijonen, a., nöjd, a., rousi, h. & kotilainen, a. 2008: marine landscapes and benthic habitats in the archipelago sea. balance interim report 31. 53 pp. https://balance-eu.org/publications/index.html (accessed november 2021). rinne, h. et al. 2014: predicting the occurrence of rocky reefs in a heterogeneous archipelago area with limited data. estuarine coastal and shelf science 138, 90–100. serrano, e. & ruiz-flaño, p. 2007: geodiversity: a theoretical and applied concept. geographica helvetica 62, 140–147. stroeven, a.p. et al. 2016: deglaciation of fennoscandia. quaternary science reviews 147, 91–121. https://doi.org/10.1016/j.quascirev.2015.09.016 tukiainen, h. & hjort, j. 2021: maisematason geodiversiteetti suomessa (landscape-scale geodiversity in finland). terra 133(2), 55–76. https:// doi.org/10.30677/terra.99435 virtanen, e.a. et al. 2018: evaluation, gap analysis, and potential expansion of the finnish marine protected area network. frontiers in marine science 5, 402. https://doi.org/10.3389/fmars.2018.00402 virtasalo, j.j. et al. 2007: late-glacial and post-glacial deposition in a large, low relief, epicontinental basin: the northern baltic sea. sedimentology 54(6), 1323–1344. https://doi.org/10.1111/ j.1365-3091.2007.00883.x walbridge, s. et al. 2018: unified geomorphological analysis workflows with benthic terrain modeler. geosciences 8, 94. https://doi. org/10.3390/geosciences8030094 https://doi.org/10.34194/geusb.v52.8317 http://www.geusbulletin.org/ https://doi.org/10.3390/geosciences9060254 https://doi.org/10.3390/geosciences9060254 https://doi.org/10.3390/geosciences9040182 https://balance-eu.org/publications/index.html https://doi.org/10.1016/j.quascirev.2015.09.016 https://doi.org/10.30677/terra.99435 https://doi.org/10.30677/terra.99435 https://doi.org/10.3389/fmars.2018.00402 https://doi.org/10.1111/j.1365-3091.2007.00883.x https://doi.org/10.1111/j.1365-3091.2007.00883.x https://doi.org/10.3390/geosciences8030094 https://doi.org/10.3390/geosciences8030094 quantifying seabed geodiversity of the archipelago sea, baltic sea, finland 1 introduction 2 material and methods 2.1 bedrock 2.2 substrate 2.3 seabed structures 2.4 quantifying the seabed geodiversity 3 results and discussion 4 conclusions acknowledgements additional information funding statement author contributions competing interests additional files references figures fig. 1 location and map of the archipelago sea along the south-west coast of finland in the northern baltic sea. the numbered ss lines in green refer to salpausselkä formations, which are large, marginal sedimentary ridges deposited during glacial standstills of the last deglaciation. the numbers refer to the age order of the formations: ss1 is the oldest, ss2 is of intermediate age and ss3 is the youngest. bathymetry data from emodnet bathymetry consortium (2020). marine and terrestrial bedrock data including faults are from gtk (2014). marine seabed substrate data are from emodnet geology (2021). terrestrial substrate data are from gtk (2018). coordinate system is euref fin tm35fin. fig. 2 seabed geodiversity was quantified based on a: rock type (modified from gtk 2014); b: seabed substrates (emodnet geology 2021); and c: seabed structures. d: background data showing specific geologic features (faults and salpausselkä formations; gtk 2014; emodnet bathymetry consortium 2020). fig. 3 distribution of seabed geodiversity in the archipelago sea, northern baltic sea. a: geodiversity index (no units). index values were calculated based on b: roughness, c: total variability as established from d: rock, e: substrate and f: structure variability. geodiversity index and roughness are represented by their minimum–maximum values. g: background data showing specific geologic features (faults and salpausselkä formations; gtk 2014; emodnet bathymetry consortium 2020). table table 1 geodiversity index statistics. richness/variability is the number of different physical elements in an area defined by a circle with a radius 2000 m. roughness represents the standard deviation of slope. andersen & pedersen 2025: geus bulletin 59. 8386. https://doi.org/10.34194/t407hg84 1 of 20 research article piggyback imbrications, duplex stacking and sequential superimposed deformation in the fanø bugt glaciotectonic complex, danish north sea lærke therese andersen1,2* , stig a. schack pedersen2 1department of geoscience, aarhus university, aarhus, denmark; 2near surface land and marine geology, geological survey of denmark and greenland, aarhus, denmark abstract this study provides a detailed structural analysis of a selected part of the fanø bugt glaciotectonic complex in the south-eastern part of the danish north sea. the 200 km2 study area was mapped in 3d using high-resolution, 2d multichannel seismic data. the interpretation of seismic profiles demonstrates an architecture markedly separated into a lower and an upper thrust fault level, separated by the upper décollement surface. the lower level is characterised by >15 thrust sheets, with crests that form subsurface ridges with reliefs up to 150 m, scattered over c. 15 km. the upper level is characterised by thrust sheets grouped in imbricate complexes with thrust faults connecting to the upper décollement. the structural style changes in the direction of transport, possibly related to the position of the ice-sheet margin responsible for the thrusting and changes in the properties of the basal décollement. the structural style is generally large-scale thrusting and folding, suggesting proglacial deformation. however, the hinterland of the upper thrust fault level displays heavily folded layers or a chaotic reflection pattern associated with subglacial deformation. special attention is drawn to two exceptional structural frameworks containing a hidden hill-hole pair: sf1, an imbricate thrust fault fan in a 5 km long and 2.5 km wide basin, developed above an extensional normal fault imbricate, and sf2, a frontal ramp uplifting an imbricated fan c. 90 m above the average level of thrusting. restored cross-sections demonstrate a shortening of the lower thrust fault level between 9–43% and 44–49% of the upper level across sf1 and sf2, respectively. we suggest that the glaciotectonic complex was formed proglacially due to gravity spreading in front of an ice margin. gravity gliding due to an inclined décollement surface of 0.5° and elevated porewater pressure at the décollement might also have facilitated the deformation. *correspondence: laerketherese@gmail.com received: 01 dec 2024 revised: 04 may 2025 accepted: 11 jun 2025 published: 06 nov 2025 keywords: seismic interpretation, multiple detachment levels, hill-hole pair, gravity spreading, gravity gliding, thin-skinned thrust-faulting abbreviations: dwtfb: deepwater thrust fault belts itfl: intermediate trust fault level ltfl: lower thrust fault level utfl: upper thrust fault level sf1: structural framework 1 sf2: structural framework 2 s.p.: shot point twt: two-way traveltime geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: kristian svennevig (geus, denmark) reviewed by: david colin tanner (liag institute for applied geophysics, germany) and one anonymous reviewer funding: see page 18 competing interests: see page 18 additional files: none provided 1. introduction glaciotectonic structures are, together with erosional and depositional features, the primary evidence of former glaciation (aber & ber 2007). glaciotectonic complexes provide a record of the nature of the deformation made by various ice advances that formed the surface morphology. in a glacial setting, gravity-spreading caused by the load of the ice sheet and frictional drag caused by the motion of the ice sheet are the driving mechanisms for the deformation of the substratum (pedersen 1987, 1996, 2000; aber et al. 1989; hart 1994; boulton & caban 1995; klint & pedersen 1995). the resulting deformation creates various structural styles in relation to the position of the ice-sheet margin. structural styles that appear in the proglacial zone differ from those formed by shearing at the sole in the subglacial zone (e.g. hart & boulton 1991; kupetz 2001; andersen et  al. 2005; pedersen 2014). in the proglacial zone, deformation is the result of longitudinal compressive stress in front of an ice sheet (e.g. aber et al. 1989; boulton & caban 1995; andersen et al. 2005). the structural styles of this zone are associated https://doi.org/10.34194/t407hg84 https://orcid.org/0000-0002-4296-2510 https://orcid.org/0000-0002-7867-5118 https://creativecommons.org/licenses/by/4.0/deed.en andersen & pedersen 2025: geus bulletin 59. 8386. https://doi.org/10.34194/t407hg84 2 of 20 geusbulletin.org with both ductile deformation (i.e. folding) and brittle deformation (i.e. thrusting), depending upon the rheology and competence of the sediments and the behaviour of the ice sheet (hart & boulton 1991; van der wateren 1995). considering the development of fold and thrust structures in glaciotectonics, it is at present regarded as being the result of fault-propagation folding (brandes & le heron 2010). in the subglacial zone, deformation is the result of frictional drag beneath the ice sheet due to the ice-sheet movement and extension caused by its load (moran 1971; hart & boulton 1991; boulton & caban 1995; van der wateren 1995; pedersen 1996, 2000; andersen et al. 2005). the resulting structural styles include folded, often refolded, pervasively sheared or chaotically mixed sediments (aber 1985), boudinage structures (berthelsen 1979) and listric normal fault imbricates (croot 1987). high porewater pressure below the snout of the ice sheet may reduce the effective load of the ice sheet to zero. this will force the proglacial compressional zone to move up-glacier to the point where the load of the ice sheet is being exerted on the subsurface again. hence, compression can occur below the ice-sheet margin (boulton & caban 1995; andersen et al. 2005). the existence of a décollement horizon at a suitable depth in the subsurface sediments is of great importance for the development of glaciotectonic structures (e.g. van der wateren 1985; huuse & lykke-andersen 2000a; andersen et al. 2005). the rheology of the décollement surface and circumstances that can alter it, such as permafrost or high porewater pressure, determine the efficiency of the décollement surface (boulton & caban 1995; boulton et  al. 1999). porewater flow and the derived elevated groundwater pressure facilitate the formation of décollement surfaces in a glacial setting (mathews & macay 1960; boulton & caban 1995; boulton et al. 1999). a dipping décollement surface may induce the mechanism of gravity gliding and contribute to deformation. structural styles such as a linkage of up-dip extension with down-dip contraction via a detachment zone is usually associated with the mechanism of gravity gliding (morley et al. 2011). according to hubbert & rubey (1959), an inclination of 0.5° is enough to facilitate gravity gliding. glaciotectonic deformation is like other compressional crustal disturbances, the only significant differences being temporal and spatial scales. in general, the size of the structures and depth of deformation are related to the size of the advancing load and the time interval during which the load was effective in reducing shear strength of the deformed material. displacement above a décollement with elevated hydrostatic pressure is essential in glaciotectonic deformation as well as in other tectonic settings (aber & ber 2007). hence, the thrust fault architecture of proglacial glaciotectonic structures is analogous to tectonic structures in thin-skinned fold and thrust belts that appear in orogenic foreland basins, at accretionary wedges or in the toes of large deltas and landslides. consequently, the same structural analysis techniques (dahlstrom 1969; hossack 1979; boyer & elliott 1982; butler 1987; de paor 1988; van der pluijm & marshak 2004) can be applied to interpret glaciotectonic structures (e.g. pedersen 1987, 1996, 2005, 2014; hart 1990; klint & pedersen 1995; van der wateren 1995; boulton et al. 1999; phillips et al. 2018). offshore mapping of the subsurface is often carried out using seismic data. applying this method to the investigation of glaciotectonic thrust fault complexes has advantages and disadvantages. often the seismic data can reveal the remains of the entire glaciotectonic complex from base to top. several profiles across the glaciotectonic complex give a good basis for performing a 3d geometric analysis and describing the internal architecture and structural styles of the thrust structures (e.g. høyer et al. 2013; pedersen & boldreel 2017; vaughanhirsch & phillips 2017; phillips et  al. 2018; winsemann et al. 2020; lohrberg et al. 2022). however, the seismic method has its limitations. the horizontal and vertical resolutions of seismic data are primarily controlled by the fresnel zone and the bandwidth of the seismic source, respectively. furthermore, steeply dipping layers or faults may not be accurately resolved, as reflections from high-angle features may fall outside the recording aperture (yilmaz 1987). hence, features smaller than the seismic resolution, and nearly vertical layers or steeply dipping faults recognised in many of the onshore coastal profiles of glaciotectonic complexes (rosenkrantz 1944; steinicke 1972; pedersen 2000, 2005; burke et al. 2009; phillips et  al. 2018; gehrmann et  al. 2019) may not be detected or imaged correctly in the seismic profiles. this study is focused on an area of 200 km2 of the fanø bugt glaciotectonic complex in the south-eastern danish north sea (fig. 1). the aim is: (1) to use high-resolution marine multichannel reflection seismic data to map and describe the detailed geometry, architecture and structural styles of the deformation framework present in the investigated area; (2) to conduct a structural analysis on two exceptional structural ‘sub-complexes’, named the structural framework 1 (sf1), situated in the centre of the area, and the structural framework 2 (sf2), situated at the mid-northern boundary of the investigated area (fig. 1). from these two ‘sub–complexes’, the structural analysis is directed towards the development of a kinematic model that supports the structural evolution of the investigated part of the fanø bugt glaciotectonic complex; (3) to discuss and suggest the possible formation mechanisms related to the mapped architecture and structural styles of the deformations. https://doi.org/10.34194/t407hg84 andersen & pedersen 2025: geus bulletin 59. 8386. https://doi.org/10.34194/t407hg84 3 of 20 geusbulletin.org 2. geological setting the fanø bugt glaciotectonic complex is situated in the south-eastern part of the danish north sea, 25 km off the west coast of south-west jylland, denmark, and covers an area of more than 600 km2 (fig. 1). the western border of the thrust complex can be mapped within the extent of the seismic data, but towards the east, northeast and south-east, the seismic sections reveal that the thrust complex extends farther towards the coast of jylland (andersen 2004). this study concerns only a minor part of the fanø bugt glaciotectonic complex (200 km2), where the seismic lines are spaced close enough to map details of individual structures. the successions affected by the glaciotectonic deformation consist of neogene silts and clays of late burdigalian and younger age (fig. 2). these sediments were deposited as a progressing accretion sourced from onshore areas of southern norway and sweden (huuse 2002). during the early miocene, sedimentation in the study area was dominantly shallow marine and terrestrial (rasmussen 1996, 2004; rasmussen et al. 2010). a regional transgression in the middle miocene changed the depositional environment to a fully marine shelf environment, which led to the deposition of the finegrained marine sediments of the hodde and gram formations (friis 1995; huuse 2002; rasmussen 2004; rasmussen et  al. 2010). the youngest neogene sediments are shallow marine or brackish-water sediments of late miocene or early pliocene age (berthelsen & kristoffersen 1974; fig. 2). the translated, faulted and folded sediments in the fanø bugt glaciotectonic complex detach on a c. 0.5° westward-inclined basal décollement surface located in the lower miocene succession (andersen 2004). the surface, which in part constitutes the basal décollement surface, displays a regional dip towards the west and south-west of the cenozoic succession in the eastern danish north sea (huuse et al. 2001). the main cause for this is the post-danian subsidence centred on the central graben and late palaeogene to recent uplift of fennoscandia (jordt et al. 1995; japsen 1998; michelsen et al. 1999). several other thrust complexes in the eastern north sea detach in miocene sediments (huuse & lykke-andersen 2000a; lohrberg et al. 2022). a common feature for several of these complexes – including the fanø bugt glaciotectonic complex – is that their direction of thrusting is almost parallel to the dip of the décollement surface (huuse & lykke-andersen 2000a). mid-latitude ice sheets transgressed parts of the north sea and melted back several times during the elsterian and the saalian glaciations, whereas the weichselian ice sheets only reached the northern part of the north sea (ehlers 1990; houmark-nielsen 2011). a succession of tills from the elsterian and saalian glaciations and fine-grained interglacial sediments from the holsteinian and the eemian succeeded the neogene deposits (knudsen & penny 1987; sha et al. 1991). quaternary buried valleys cutting both the neogene and quaternary successions are recognised in several places in the north sea (salomonsen 1995; huuse & lykke-andersen 2000b; andersen et al. 2013; steward et al. 2013; prins & andersen 2019; lohrberg et al. 2020; ottesen et al. 2020; wenau & alves 2020) as well as onshore denmark (andersen et al. 2013; sandersen & jørgensen 2017). fig. 1 location in the south-east danish north sea of the fanø bugt glaciotectonic complex and the seismic lines used for interpretation. the 10 × 20 km2 study area is marked with a red rectangle. locations of the two structural frameworks sf1 and sf2 are marked with purple and blue rectangles, respectively. locations of seismic lines shown in figs 3–6 are outlined in yellow. jylland, denmark 0 25 km 55°00’ 55°30’ 55°30’7°30’ 0°e 20°e 60°n dk de se no uk germany study area deep well sf2 n sf1 fi g. 5 fi g. 5 fi g. 6 fig. 4 fig. 3 s-1 fanø røm ø sy lt https://doi.org/10.34194/t407hg84 andersen & pedersen 2025: geus bulletin 59. 8386. https://doi.org/10.34194/t407hg84 4 of 20 geusbulletin.org the upper boundary of the deformed succession in the fanø bugt glaciotectonic complex is interpreted to be a saalian glacial unconformity. this erosional surface forms a hummocky and even hilly landscape with small erosional basins (larsen & andersen 2005). the deformation of the glaciotectonic complex is not directly dateable, as no adequate borehole information is available in the fanø bugt area. an elsterian or early saalian age has been suggested (huuse & lykke-andersen 2000a). kinetostratigraphic correlation with onshore glaciotectonic deformations of similar size, style and transport direction indicates that the fanø bugt glaciotectonic complex formed during a stage of westward movement of the late saalian (warthe) ice sheet (andersen 2004; houmark-nielsen 2007; jørgensen et  al. 2012). a pre-elsterian or early elsterian age of formation has recently been suggested by winsemann et al. (2020), based on a correlation with the glaciotectonic complex at heligoland in the south-eastern german north sea. 3. data and methods 3.1. seismic data acquisition and parameters the university of aarhus and the geological survey of denmark and greenland (geus) acquired the high-resolution marine seismic data on scientific cruises in the north sea during the years 1999–2001 with the danish navy ship, flyvefisken. the seismic data from these surveys, named fl99, fl00 and fl01, were collected to map the area of the fanø bugt glaciotectonic complex (andersen 2004). the surveys make up a dense grid of 1065 km long, high-resolution, multichannel-reflection seismic lines, distributed over an area of c. 10 × 20 km2 (fig. 1). the grid spacing is c. 250 m between ne quat. holocene pleistocene pl io . l e piacenzian zanclean messinian tortonianl serravallian m m io ce n e n eo g en e e 10 15 20 25 30 chattian langhian burdigalian aquitanian l e rupeliano li g o ce n e 5 pa la eo g en e pe ri o d sw ep o ch m a ag e se is m ic re fl ec ti o n lithostratigraphy marine silt and clay hiatus brackish water silt and clay quaternary erosion fluvial and marine sand gs coal branden fm. ? linde clay viborg fm. odderup fm. arnum fm. gram fm. hodde fm. bastrup sand basal décollement surface re�ections on the seismic pro�les correlated with s-1 well vejle fjord fm. ribe fm.?? fig. 2 stratigraphic diagram (modified from rasmussen 2004). the succession involved in the thrusting consists of fine-grained sediments (the south-western part of the diagram) of late burdigalian and younger age. the basal décollement surface (black line) developed in the fine-grained sediments of the arnum fm (early miocene). the upper décollement surface developed in the fine-grained sediments of the hodde fm (middle miocene), between the orange and purple reflectors. https://doi.org/10.34194/t407hg84 andersen & pedersen 2025: geus bulletin 59. 8386. https://doi.org/10.34194/t407hg84 5 of 20 geusbulletin.org approximate dip lines and c. 1500 m between approximate strike lines. in addition, a few lines from the surveys da95 and gr98 (huuse 2000), which crossed the area, were used, and a few conventional oil-seismic data (rtd–re94 and np85n) were used for the correlation of the interpreted horizons to the well s-1 about 70 km west of the study area (andersen 2004). the signal bandwidth of the migrated seismic data is 50–180 hz, and the vertical penetration is 2000 ms twoway traveltime (twt), corresponding roughly to 2000 m (clausen & huuse 1999). the seismic data acquisition was limited towards land by the 10 m water-depth contour, selected as a safety precaution for equipment and ship. the vertical and horizontal resolutions are approximately 5 m and 20 m, respectively. data quality is generally good; however, interpretation of the top 50–100 ms twt is often hampered by water-layer multiples. minor mis-ties between individual surveys and mis-ties due to migration problems in the neighbourhood of steeply dipping faults and associated tilted layers were identified and corrected prior to interpretation. the limited resolution and challenges in imaging steep dips may explain the presence of large areas of chaotic reflection patterns (e.g. fig. 3, shot point, s.p., 0–730, top half of the profile). it is likely that the structures in this area were either smaller than the seismic resolution (5 m vertical and 20 m horizontal) or that layers or faults were steeply dipping. however, the nature of the deformation can also contribute to obtain a chaotic reflection pattern in areas where structural and sedimentary architecture is particularly complex. 3.2. interpretation, maps and plots the seismic data were interpreted digitally using landmark’s openworks, seisworks 2d software (2002–2003). difficulties occurred when interpreting thrust horizons, as the software does not allow multiple twt values for a single horizon at the same shot point (s.p.). hence, an advanced interpretation strategy was necessary. as the overlap of a thrust horizon cannot be represented by a single horizon, a new horizon was made where a twt value occurred at the same s.p. representing the deeper twt values of overlap between the adjacent thrust sheet. hence, all the upper twt values were represented by the same horizon. maps of the basal décollement (floor thrust), the upper décollement (roof thrust) and a structural map of an internal reflection in the lower thrust level were drawn using landmark’s openworks, z-map+ software. the horizons data were imported from seisworks and were gridded using the least squares algorithm, with a grid increment on 45 m. smoothing was performed before contouring with an interval of 10 ms and 20 ms, respectively. the dip of the basal décollement surface was calculated using a velocity of 1850 m/s and basic trigonometry by the following equation: tan(a) = a/b (eq. 1) where a = dip angle, b = distance between contours perpendicular to the strike of contours measured on the map. a is defined in eq. (2) as follows: fig. 3 seismic section (fl01–50) showing a duplex with imbricated thrust sheet in two levels. the lower thrust fault level (ltfl) is detaching on the basal décollement surface acting as a floor thrust. the utfl is detaching on the upper décollement surface, acting as a roof thrust for the ltfl and a floor thrust for utfl. ltfl and utfl can be divided into three and two structural domains, respectively, showing different structural styles (see section 4.3 and 4.5 and fig. 8). the seismic section is an approximate dip line to sf1, showing piggyback imbricates in a basin (s.p. 725–900) behind a huge hanging wall anticline (s.p. 900–1050). seismic section shown with 5x vertical exaggeration. interpretations shown below the seismic section with no exaggeration. here, the inclination of the thrust faults ramps is measured. crossing seismic lines are marked on the seismic section with a figure number, see figs 1, 7 and 8 for location of this cross-section. s.p.: shot point. twt: two-way travel time. fig. 5fig. 6 fig. 5fig. 6 vertical exaggeration 5x no exaggeration inclination of thrust ramps 34° 32° 35° 25° 25° 15° 15° 15° 15°23° 23° 20° 1000 m ew twt (ms) utfl sf1 distal domain utfl proximal domain utfl proximal domain ltfldistal domain ltfl extension extension extension ? intermediate domain ltfl hangingwall anticline strongly sheared sedimentspiggyback imbricates in the basin of sf1 ltfl s.p. s.p. 1002003004005006007008009001000110012001300 pull up 1002003004005006007008009001000110012001300 0 500 0 500 0 500 0 500 https://doi.org/10.34194/t407hg84 andersen & pedersen 2025: geus bulletin 59. 8386. https://doi.org/10.34194/t407hg84 6 of 20 geusbulletin.org a = ((contourtwtmax – contourtwtmin)/2) × velocity (eq. 2) the velocity used here is an average velocity of the post chalk group, based on nielsen & japsen (1991). a velocity range of 1700–2000 m/s was considered for uncertainty, resulting in a variation in dip angle of approximately ±0.03°. plots of the seismic data and interpretation were produced for illustrations. the vertical scale was calculated using a velocity of 1850 m/s, while the horizontal scale was determined by multiplying the number of traces by the distance between traces. the seismic data plots were produced with a 5x vertical exaggeration. a version without vertical exaggeration, displaying only the interpretation, is presented below the seismic section. this version was created by reducing the vertical scale by a factor of five while preserving the horizontal scale. 3.3. structural analysis a structural analysis was conducted based on the mapping and description of horizon topography, thrust fault orientation and geometry of thrust sheets. restoration of sf1 and sf2 was performed to give a rough estimate of the minimum amount of shortening and to reconstruct the deformational history of sf1 and sf2 in a tentative kinematic model. the concept for the restoration of the cross-section was based on the method for balanced cross-sections (dahlstrom 1969; de paor 1988; wilkerson & dicken 2001; van der pluijm & marshak 2004). this method assumes constant layer thickness, constant along-strike displacement along faults and plane-strain deformation. we realise that these assumptions cannot all be fulfilled in a glaciotectonic context, and that the limitations of the seismic method have implications for the result. when performing the structural balancing, it is essential to restore all strains and displacements in the reverse order in which they were applied (butler 1987). hence, the restoration was done stepwise, starting with the youngest, westernmost faults in the lower thrust fault level (ltfl) towards the oldest, easternmost faults. the same procedure was performed for the structures in the upper thrust fault level (utfl), as it was assumed that the utfl was deformed before the ltfl (see section 4.6 for more details). we performed a tentative 2d line-length balancing of the structural interpretation of the two seismic cross-sections, approximate dip lines to sf1 and sf2 (figs 3 and 4) on the interpreted reflections, the purple reflector in the utfl and green reflector in ltfl (results in sections 4.8 and 4.9). the starting point for the kinematic model was the deformed sections (the interpreted seismic sections) illustrated in the final step of the restoration. from here, fig. 5 fig. 5 vertical exaggeration 5x no exaggeration 500 m ew twt (ms) itfl ? s.p. s.p. 0 500 0 500 0 500 0 500 100200300400500600700800 100200300400500600700800 utfl hanging wall anticline ltfl distal domain ltfl intermediate domain ltfl basal décollement upper décollement intermediate décollement sf2distal domain utfl piggyback imbricates on duplex fig. 4 seismic section (fl01–54) showing an approximate dip profile across the structural framework sf2, showing piggyback imbrication of the sedimentary succession of ltfl repeated and uplifted more than 100 ms (c. 90 m; see also sf2 in fig. 8) on a frontal ramp and a flat acting as an intermediate décollement surface for itfl. west of sf2 is the general duplex of the area with large-scale thrust structures (some more than 1000 m long) in ltfl and utfl translating across the basal décollement and upper décollement, respectively. hanging wall anticlines in ltfl is superimposed in the upper décollement as folding of the surface. small-scale structures can be seen in the hinterland part of the cross-section but cannot be outlined accurately. seismic section shown with 5x vertical exaggeration. interpretations shown below the seismic section with no exaggeration. crossing seismic lines are marked on the seismic section with a figure number. see figs 1, 7 and 8 for location of this cross-section. s.p.: shot point. twt: two-way travel time. https://doi.org/10.34194/t407hg84 andersen & pedersen 2025: geus bulletin 59. 8386. https://doi.org/10.34194/t407hg84 7 of 20 geusbulletin.org the model progresses stepwise up to step 1 – the balanced section. folding is generally regarded as fault-related; however, in one place, we tentatively interpret the deformation as ductile attributed to hydrodynamic shear deformation (see section 4.9; step 3, top of sf13). a sediment package subject to hydrodynamic shear deformation could potentially spread in multiple directions. hence, in this case, line balancing is first applied to the tilted normal fault blocks, and the area subjected to hydrodynamic shear deformation is reconstructed to fill in the rest of fault block sf13. since the restoration was performed only on one reflector in each level, the fault dip was preserved as observed in the deformed section throughout the restoration process. to calculate section shortening, a fixed pin line was placed at the front of each thrust-fault level and a loose pin in the rear. the displacement between the deformed and restored sections provides the average contraction of the upper and lower thrust fault levels, respectively. the calculations were done by measuring the distance between the front (fixed) pin and the rear (loose) pin for each thrust-fault level and for each step, relative to the original length. erosion of the top of the sections is not accounted for. hence, the calculated shortening represents a minimum estimate. an exception is the westernmost fault block of the utfl in sf1, where an eroded fault block segment is added to maintain the front pin line position, as this fault block is moved eastward during the restoration of the ltfl from step 4 to step 3 (see section 4.9). 4. results 4.1. the seismic interpretation the seismic data quality is generally good and provides many excellent images of the glaciotectonic structures in a large part of the study area. the line spacing in the w–e direction is close enough to trace the largest faults from line to line, providing a good opportunity to describe the geometry, the internal architecture and structural styles of this part of the fanø bugt glaciotectonic complex. the interpretation is based on the recognition of distinct reflection patterns; equal thickness of undeformed layers across faults was used in areas where the identification of the reflection pattern was difficult. ramps of thrust faults or normal faults were mapped where an abrupt cut of reflections occurs in the vertical plane. thrust-fault flats or décollement surfaces were mapped where more or less horizontal discontinuities existed. these boundaries appeared where undeformed sediments were overlain by deformed sediments, or where different structural patterns were recognised across a horizontal or weakly inclined surface. five internal reflections, including two décollement surfaces and numerous thrust faults and normal faults, were interpreted. the following description of the interpreted and mapped surfaces and fault is exemplified on four seismic cross-sections: two approximate dip sections (figs 3 and 4) and two sections oblique to strike (figs 5 and 6) cross the centre of sf1 and sf2, respectively (see fig. 1 for location). the internal architecture fig. 5 seismic section (fl00–27) is crossing the area oblique to strike of the structural elements and across both sf1 and sf2. the cross-section is showing an apparent ‘spreading’ of the transport direction of thrust sheets between thrust fault levels. the apparent direction of thrust sheets of utfl is towards south (s.p. 570–1300). at sf2, the lower succession is repeated in itfl; here, the apparent thrust direction is towards north (s.p. 1500–1800). the apparent direction of thrusting of ltfl is towards south, ending at s.p. 900. south of the study area, thrusting only appears in the utfl. normal faults in the ltfl are seen at sf1 dipping opposite the thrust fault. the utfl repeats across sf1 (see purple reflector), compare also with dip-section across sf1 (fig. 3). the true thrust direction can be inferred from fig. 8. seismic section shown with x5 vertical exaggeration. interpretations shown below the seismic section with no exaggeration. crossing seismic lines are marked on the seismic section with a figure number. see figs 1, 7 and 8 for location of this cross-section. s.p.: shot point. twt: two-way travel time. twt (ms) s.p. s.p. no exaggeration vertical exaggeration 5x ns fig. 3 fig. 4 1000 m 0 500 0 500 0 500 0 500 210020001900180017001600150014001300120011001000400 500 600 700 800 900 210020001900180017001600150014001300120011001000400 500 600 700 800 900 sf1 sf2 study area piggyback imbricates/ itfl folding of utfl hanging wall anticlines complex area between sf1 and sf2 on fig.9 basal décollement https://doi.org/10.34194/t407hg84 andersen & pedersen 2025: geus bulletin 59. 8386. https://doi.org/10.34194/t407hg84 8 of 20 geusbulletin.org and structural styles seen in these figures are described in the figure caption. 4.2. the mapped surfaces the lower boundary of the glaciotectonic complex is the basal décollement surface. this surface separates the undeformed layers below from the deformed layers above, see the seismic lines figs 3–6. the surface that constitutes the basal décollement surface within the study area has been mapped in a larger area, and a twt map has been generated (fig. 7). this map is described in section 4.4. the layers just above the basal décollement surface have a distinct reflection pattern of several parallel (when not disturbed) very strong reflections. the top reflection of this unit is marked with green in the seismic cross-sections (figs 3–6). it is possible to follow the green reflector in most of the study area, and a structural twt map shows the topography of this surface and the thrust fault trends of the ltfl (fig. 8). a description of this map is given in section 4.5. the reflection pattern above the green reflector also consists of parallel reflections, although less intense. the orange reflector marks the top of three (sometimes only two) very strong reflectors in the upper part of the ltfl (figs 3–6). because of disturbance due to deformation, it has only been possible to map this surface in parts of the study area, and no resulting topographic map has been generated. the roof thrust of the ltfl appears above the orange reflector. we refer to it as the upper décollement surface (figs 3–6). above this surface, there is the utfl. the upper décollement surface sometimes displays a very strong reflection, and in other parts, it was mapped where a marked change in structural styles appears above and below the surface. a perspective view of this décollement surface is seen in fig. 9 and described in section 4.6. in the utfl, only one reflector was outlined, which, in the seismic cross-sections, is displayed in purple (figs 3–6). this surface is the reflection from a series of parallel-layered beds above the upper décollement surface. these are easily recognised in the undeformed layers west of the thrust front (figs 3–6). the interpretation of the purple reflector was hampered by multiples and deformation of the layers, making it difficult to recognise in large parts. hence, no resulting topographic map was generated of this reflector. above this surface, the reflection pattern is semi-parallel and generally noisy and often distorted by multiples (figs 3–6). the westward continuation of the basal décollement surface is stratigraphically located in the lower miocene succession (andersen 2004), and the green, orange and purple reflectors are in the middle miocene succession, when correlated with the s-1 well, situated c. 70 km nw of the fanø bugt glaciotectonic complex (fig. 2). the ltfl is representing the early/middle miocene marine arnum formation with possible intercalations of vertical exaggeration 5x no exaggeration fig. 3 fig. 3 twt (ms) s n 500 m s.p. s.p. 0 500 0 500 0 500 0 500 2100 2000 1900 1800220023002500 24002600 2100 2000 1900 1800220023002500 24002600 sf1 study area folded utfl upper décollement surfaceupper décollement surface hanging wall anticlines extensional faults in ltflbasal décollement fig. 6 seismic section (fl01–19) is crossing the area oblique to the strike of structural elements and across sf1 at s.p. 1800–2200. ltfl and utfl show apparent thrusting towards south. thrusting of utfl is continuing further towards south (out of the study area) than deformation of ltfl. in the northern part of the profile, extensional faulting is seen in ltfl below the depression of the upper décollement surface. utfl is repeated in the basin of sf1 (s.p. 1880–2030, see purple reflections). the true thrust direction can be inferred from fig. 8. seismic section shown with 5x vertical exaggeration. interpretations shown below the seismic section with no exaggeration. crossing seismic lines are marked on the seismic section with a figure number. see figs 1, 7 and 8 for location of this cross-section. s.p.: shot point. twt: two-way travel time. https://doi.org/10.34194/t407hg84 andersen & pedersen 2025: geus bulletin 59. 8386. https://doi.org/10.34194/t407hg84 9 of 20 geusbulletin.org the sandy bastrup and odderup formations, whereas the utfl constitutes the middle miocene, fully marine, clayey hodde and gram formations (fig. 2). the structural styles of the utfl are generally more ductile than the deformation in the ltfl, which is interpreted to be due to the different rheology of the sediments. thrust-fault trends in the utfl were mapped from the seismic cross-sections, which provide the transport direction of the thrust sheets (fig. 10 and section 4.7). the structures of the utfl are eroded or truncated at the top by a possible late saalian unconformity (larsen & andersen 2005) or the seafloor (fig. 3, s.p. 1000–1200). 4.3. architecture and structural styles the thrust fault architecture of the study site indicates that the overall geometry is a large-scale duplex with imbricate structures bounded by a floor thrust (the basal décollement) and a roof thrust (the upper décollement). n6140000 8˚ e e4 40 00 0 e4 50 00 0 e4 60 00 0 n6130000 55º15 55º20 n6120000 n6110000 8˚ e e4 40 00 0 e4 50 00 0 n6130000 55º15 55º1055º10 n6120000 n6110000 300 400 55º20 fi g. 5 fi g. 6 fig. 4 fig. 3 depth in twt (ms bsl.) 0 5 km projection ed50, utm 32 jylland, denmark røm ø fanø study area sf2 sf1 deformation front normal fault dip direction area with pull up e�ect fig. 7 a two-way time-contour map of the basal décollement surface and its continuation beyond the western limit of ltfl deformation. the deformation front shows that the maximal western extent of the fanø bugt glaciotectonic complex is situated in the study area. the contour curves of the surface show a marked change in dip and dip direction at the border of the study area (see section 4.4 for further details). a pull-up of the surface below the basin of sf1 is marked with a black dashed line. locations of sf1 and sf2 are shown. a small normal fault and a depression of the surface are seen in the southern part of the study area. seismic sections shown in figs 3–6 are marked with black lines. fig. 8 a structural two-way time (twt) contour map showing the strike of the main structures at the level of an internal reflection in ltfl and itfl (green reflector on figs 3–6). to enhance the overview, the map only shows the larger faults. thrust faults are marked with red, and normal faults are marked with black lines. the change of strike of faults reveals different structural domains. in the eastern proximal domain, the faults strike 160°, in the intermediate domain, the strike of faults is 155° and in the western distal domain, the strike is 170°–180°. the three structural domains are showing different structural styles (see section 4.5 and fig. 3). the trend of the faults is turning in the southern part of the area from west south-west to south-west following the same depression in the surface as is seen in the basal décollement surface (fig. 7). in the area of sf2, the mapped reflector is elevated c 100 ms (c 90 m) on a huge frontal ramp and flat and display the structural trend of itfl. the remaining parts of the map show the structural trend of ltfl. the structural trend of itfl is towards the north-west, hence different to the general structural trend of ltfl, which is towards the west south-west. sf1 displays two huge hanging wall anticlines in front of an extensional basin/depression. the location of sf1 and sf2 is market with rectangles. the seismic sections shown in figs 3–6 are marked with black lines. 0 5 km projection ed50, utm 32 n6140000 8˚ e e4 40 00 0 e4 50 00 0 e4 60 00 0 n6130000 55º15 55º20 n6120000 n6110000 8˚ e e4 40 00 0 e4 50 00 0 n6130000 55º15 55º1055º10 n6120000 strike 160° strike 160° strike 155° strike 155° strike 170-180° strike 170-180° n6110000 fi g . 5 fi g. 6 fig. 4 fig.3 100 150 200 250 300 350 jylland, denmark røm ø fanø distal domain intermediate domain proximal domain study area sf2 sf1 thrust fault extensional fault strike of faults location of �gs.3-6 https://doi.org/10.34194/t407hg84 andersen & pedersen 2025: geus bulletin 59. 8386. https://doi.org/10.34194/t407hg84 10 of 20 geusbulletin.org the ltfl shows three deformation domains, characterised by different structural styles (figs 3 and 8) as follows: 1. small imbricated thrust sheets are seen in the proximal domain to the east (fig. 3). 2. larger thrust sheets with tailing tilted normal faults are seen in the intermediate domain (fig. 3). 3. the distal domain constitutes large thrust sheets with hanging-wall anticlines. close to the deformation front in the west, back thrusts and pop-up structures dominate the distal domain (figs 3 and 4). a more thorough account is presented together with the map of the ltfl (fig. 8) in section 4.5. the succession above the upper décollement surface, the utfl, displays two different deformation domains: (1) a distal domain with an imbricate system of piggyback thrust structures (figs 3 and 4); (2) a proximal domain with a diffuse or chaotic reflection pattern (fig. 3), interpreted to resemble pervasively sheared or chaotically mixed sediments, possibly involving boudinage structures, fluid flow structures and hydrodynamic brecciation. in the central part of the study area, the structural framework sf1 dominates the seismic images (see dip section in fig. 3, s.p. 760–1020, sections oblique to strike in fig. 5, s.p. 1100–1300 and fig. 6, s.p. 1800–2200), showing a piggyback imbricate fan of the utfl in a depression 5 km long and 2.5 km wide. below the piggyback imbricate fan, several partly rotated normal faults are present. in front of the depression, a huge thrust sheet forms a hanging-wall anticline repeating the ltfl in the narrow crest of the anticline (fig. 3, s.p. 900–1050 and fig. 6, s.p. 2100–2200). in the mid-northern part of the study area, another spectacular structural framework is displayed in the seismic sections sf2 (see dip section in fig. 4, s.p. 0–250 and the section oblique to strike in fig. 5, s.p. 1500– 1800). sf2 consists of a ramp and a flat, along which an imbricate fan of thrust sheets is translated westwards and leaves a tailing depression to the east. the imbricate fan formed by ramp collapse of a thrust segment derived from the ltfl and raised c. 90 m above the average level of thrusting in the ltfl. this flat acts as an intermediate décollement surface, and, hence, the imbricated fan represents an intermediate trust fault level (itfl), situated between an intermediate décollement surface and the upper décollement surface (fig. 4, s.p. 0–250). the imbricate fan repeats the sedimentary succession of the ltfl in an area of approximately 8.5 km2, which now comprises a structural high (see sf2 in fig. 8). the eastern part of sf2 is seen as a depression at the level of the green reflection, which fig. 9 3d map of the upper décollement surface with a 2.5x exaggeration gives a perspective view of the hilly landscape with ridges above hanging wall anticlines and depressions above extensional faulting in the ltfl. both sf1 and sf2 comprise hill-hole pairs at this surface level, although the internal structural framework is very different (figs 3 and 4). sf1 and sf2 are connected by a low-lying area with a complex reflection pattern; hence, the interpretation of this area is uncertain (see section 4.6). there is a remarkably large coherent ridge just proximal sf1 and sf2. sf1 sf2 complex area/tear fault? 25 0 20 0 15 0 10 0 502.5 km depth in twt (ms) n fig. 10 map comparing the trends of the faults of the utfl (black) compared with those of the ltfl (grey), see section 4.7. the structural trend of the utfl generally follows that of the ltfl in a smooth manner, for example, across sf2, and the structures of the utfl connect the two fault directions of ltfl and itfl in smoothly curving faults. in the south, the bending faults in the ltfl seem to affect the trend of the utfl much more gently. jylland, denmark røm ø fanø 0 5 km projection ed50, utm 32 8˚ e e4 40 00 0 e4 50 00 0 n6130000 55º15 55º20 n6120000 8˚ e e4 40 00 0 e4 50 00 0 55º15 55º1055º10 n6120000 sf2 sf1 fault utfl fault ltfl https://doi.org/10.34194/t407hg84 andersen & pedersen 2025: geus bulletin 59. 8386. https://doi.org/10.34194/t407hg84 11 of 20 geusbulletin.org is partly due to extensional faulting of the ltfl (fig. 8) and partly caused by the removal of material due to the thrust-sheet translation described earlier (figs 4 and 8). 4.4. the basal décollement surface a time-contour map of the basal décollement surface is presented in fig. 7. north of the study area and in the north-eastern corner of the study area, the surface dips towards southwest. inside the study area, the dip is gentler towards west-south-west, changing to south-south-west in the south. south of the study area, the surface dips towards north-west and south-west due to a depression in the surface (fig. 7). the strike of the contours changes from c. 155° to 160° at the 280 ms twt curve in the hinterland to c. 170° at the 340 ms twt curve in the foreland. furthermore, a change in dip angle characterises the transition from the hinterland with low dip (c. 0.3°) in the east to a slightly steeper dip towards the foreland (c. 0.4°–0.6°). the steepest dip is just east of sf1. in the south-west, the contours change direction again due to a depression of the surface in the south-western corner, and the dip increases to c. 1.5° (fig. 7). this depression and the small normal fault in the southern part of the study area are also seen in the top chalk surface (clausen & huuse 1999). the contour curve at 320 ms twt of the basal décollement surface makes an upward bulge just at the base of sf1, marked with an elongate circle of black dots in fig. 7. this is presumably a ‘pull-up’ effect caused by higher seismic velocities in the piggyback-imbricate thrust sheets filling the basin (fig. 3, s.p. 850-925). in the surrounding area, the décollement surface is relatively smooth, with no obvious indications of a real bulge on the décollement surface. the existence of velocity pull-up is supported by the seismic image in fig. 3 (see basal décollement, s.p. 850–925). the deformation front of the ltfl (fig. 7) marks the transition from the deformed frontal part of the complex to the undeformed foreland towards west. the deformation front reaches its maximal western extent in the study area (fig. 7). towards the south-west, the deformations reach their maximum depth and coincide approximately with the 400 ms twt contour curve of the basal décollement (fig. 7). this may indicate that c. 360 m (c. 400 ms twt) is the maximum deformation depth in the fanø bugt glaciotectonic complex. 4.5. the structural trend of the lower and intermediate thrust fault levels a structural two-way time-contour map was generated (fig. 8) to investigate a distinct internal surface in the lower part of the complex (indicated by the top level of the green reflector in figs 3–6). the map shows both the topography of the surface and the geometry, dimension and transport direction of the structures of the ltfl translated on the basal décollement surface. furthermore, fig. 8 shows the structures of the intermediate thrust fault level (itfl) translated on the intermediate décollement surface at the location of sf2. thrust faults and back thrusts are displayed by red lines and normal faults with black lines. for an overview, fig. 8 shows only 25 of the larger faults out of a total of 35 faults shown for the ltfl in fig. 3. the fault trend in the ltfl can be divided into three structural domains based on the change in strike of the fault structures (fig. 8). on the seismic sections, the change in fault strike correlates with a change in structural styles (see ltfl in fig. 3). the three structural domains are as follows: 1. the eastern proximal domain consists of a succession of relatively small imbricated thrust sheets striking 160°, with thrust faults dipping c. 15° (fig. 3, s.p. 125–335). 2. in the intermediate domain, a 2 km wide block of more or less undeformed sediments (fig. 3, s.p. 335– 480) constitutes the eastern part of the domain (fig. 8). farther west, a succession of larger thrust sheets with tailing extensional listric normal faults exists. the faults are striking 155° (fig. 8) and dipping c. 20°– 25° (fig. 3, s.p. 480–750). the significant structure sf2 is located in the central northern part of the intermediate domain (fig. 8). sections crossing this structure are shown in fig. 4 (s.p. 0–250) and fig. 5 (s.p. 1500– 1800). a correlation of a hill-hole pair is expressed by the depression east of the ‘thrust-fault high’ (fig. 8). at the ‘thrust-fault high’, the ltfl is repeated directly on top of itself with the utfl above. hence, at this place, one could argue that a third itfl is present, which, however, consists of the same sediment package as the ltfl (see e.g. fig. 4, sf2). the map (fig. 8) always expresses the upper position of the mapped green horizon, so, in areas where the horizon repeats itself, it is the itfl, which is represented in the ‘thrust-fault high’ (fig. 8). the structural trend of itfl in sf2 is towards the north-west, which is notably different to the general structural trend of ltfl, which is towards west and west-south-west (figs 5 and 8). 3. the western distal domain contains faults with a strike direction of 170°–180° (fig. 8). the significant sf1 appears in this domain (fig. 3, s.p. 750–1000). in the frontal part, sf1 comprises a more than 1000 m long thrust sheet (fig. 3, s.p. 900–1000). the inclination of thrust faults in sf1 is 25–23°, the same as for https://doi.org/10.34194/t407hg84 andersen & pedersen 2025: geus bulletin 59. 8386. https://doi.org/10.34194/t407hg84 12 of 20 geusbulletin.org the structures in the intermediate domain (fig. 3, s.p. 950–1000). however, in the map view, the strike of the faults in sf1 is the same as for the distal domain (fig. 8). in the tailing part of sf1, tilted normal faults created a depression (fig. 3, s.p. 775–900). a spatial view of this depression can be seen in figs. 8 and 9. this depression covers an area of c. 5 km × 2.5 km, where it is most pronounced. in the distal domain west of sf1, the structural style comprises relatively large thrust sheets with pop-up structures in the frontal part, dipping c. 25°–35° (fig. 3, s.p. 950–1270). these structures delineate the western boarder of the thrust complex (fig. 8). small thrusts or reverse faults, with fault planes dipping towards the west, express the thrust front towards the north-western corner of the study area (fig. 4, s.p. 700). these structures could have started as nucleation points of conjugate contractional faults (andersen et  al. 2005), which developed into pop-up structures and fore-thrust/ back-thrust pairs during the progressive compressional deformation. 4.6. the upper décollement surface the upper décollement surface is interpreted to be a decoupling unconformity separating utfl from ltfl (see figs 3–6) in most of the study area. a perspective view of the upper décollement surface (fig. 9) shows a hilly topography with variations of about 250 ms twt (c. 225 m) from low to high areas. the topography appears as a series of ridges elongated in a north-west – southeast direction, which probably represents folding introduced by superimposed sequential deformation over hanging-wall anticline crests and small extensional basins in the ltfl. the bumpy nature of the surface, especially in the east, is probably a combination of differential displacement across thrust faults in the ltfl and uncertainties in picking the reflection of the upper décollement surface, as this is not a clear and unambiguous reflection in all places (e.g. see fig. 3, s.p. 0–350). at the location of both sf1 and sf2, a hill-hole pair is outlined in the upper décollement surface (fig. 9). the glaciotectonic feature referred to as hill-hole pair is applied in the sense of aber et al. (1989). a low-lying, elongated area striking west-south-west – east-north-east is situated between the two hill-hole pairs (fig. 9). a tentative interpretation of this area could include a tear fault departing the two structures, sf1 and sf2. however, as sf1 and sf2 have different internal architectures, it is not a tear fault in the traditional sense, where the same structure can be found on each side of the tear fault (hills 1963). however, in a glaciodynamic setting, the two structures could have undergone further deformation individually after the separation. the low-lying area shows a complex reflection pattern; hence, the seismic interpretation here is uncertain, probably due to the limitations of the seismic method and the complexity of the structure (e.g. fig. 5, s.p. 1400–1500). 4.7. the structural trends in the utfl the structural trends of the thrust faults in utfl are mapped to see if there were any discrepancies between the structural trends of utfl, itfl and ltfl, which could imply stress from different directions (fig. 10). the thrust front to the west ends at the same location for the utfl and the ltfl (fig. 10). in general, the trends of the structures in utfl are notably like those of the ltfl and itfl, except for small discrepancies (fig. 10), indicating that they have been subject to similar stresses during formation. the structures of utfl can be divided into two structural domains based on fig. 3, a westerly distal domain, where imbricate thrust structures are translocated along the upper décollement surface until the eastern margin of sf1 (fig. 3, s.p. 730–1260), and an easterly proximal domain, where the sediments are strongly deformed (see fig. 3, s.p. 0–730). the thrusting of utfl continues south of the study area, where ltfl is not deformed (fig. 5, s.p. 550–880). the seismic lines in this area are, however, so far apart that the individual faults cannot be traced. strike sections in the study area show large (up to c. 4–6 km) folded thrust sheets of utfl in connection with and in between sf1 and sf2 (fig. 5, s.p. 1200–1600 and fig. 6, s.p. 2260–1875). 4.8. restoration of the deformed sections the method for performing the restoration is described in section 3.3. the method is only giving a rough estimate of the minimum shortening across the two structural frameworks sf1 and sf2. figure 11 shows the restored section across sf1 from s.p. 550–1100 (fig. 3). the restoration shows a minimum average shortening of utfl of 44%, corresponding to a lateral displacement of 3400 m and only 9% average shortening of ltfl, corresponding to a lateral displacement of 600 m. the shortening of utfl across sf1 is very strong with large displacements across each thrust fault. the displacement is generally much less in the ltfl, where extension at some points is the prevailing deformation (step 3). hence, the utfl shows more compression than ltfl in the cross-section across sf1. figure 12 represents the restoration of a section across sf2 in the mid-north of the study area from s.p. 0–440 (fig. 4). in this cross-section, the ltfl is repeated in an imbrication fan above the normal level of ltfl. the utfl is forming a duplex stack in front of the imbrication https://doi.org/10.34194/t407hg84 andersen & pedersen 2025: geus bulletin 59. 8386. https://doi.org/10.34194/t407hg84 13 of 20 geusbulletin.org fan (step 3 and step 4 in fig. 12). the average minimum shortening of the utfl is 49% at sf2, corresponding to a lateral displacement of c. 2.2 km. the shortening of ltfl is 43%, corresponding to a lateral displacement of c. 3.8 km. the restoration shows the amount of shortening in each thrust-fault level varies from place to place within the thrust-fault complex. data from the fringing areas of the study area show that the deformation continues farther east of the mapped area, towards the coast. therefore, a full account of shortening of the fanø bugt glaciotectonic complex cannot be provided with the existing data. 4.9. the structural framework and dynamic development of sf1 and sf2 the structural framework of sf1 is situated in the distal domain, although it has many similarities with the structures in the ltfl intermediate domain (fig. 3). they both involve extensional faulting and have a frontal-thrust ramp with a low dip (fig. 3). however, the strike of the faulting is like the strike of the structures of the distal domain (fig. 8). the sequential development of sf1 comprises three dynamic phases illustrated in four steps in fig. 11. fault blocks in the ltfl are coloured green and annotated sf11–sf14. the normal fault imbricates (the extensional area) are grouped in sf13. the fault blocks in the utfl are not annotated but coloured purple (fig. 11). the four steps are described as follows: 1. step 1 (fig. 11) presents the balanced cross-section of sf1, where the thrust sheets are organised into their pre-deformation position. 2. step 2 (fig. 11) illustrates the first phase of initial deformation, where the utfl is laterally displaced about 2.15 km, corresponding to 28% over the upper décollement surface, and the ltfl is compressed by 6% along the basal décollement surface. at the tailing end of the utfl, a thrust-fault imbricate fan formed, which corresponds to the 2.15 km contraction. note that the deformation of utfl occurred slightly ahead of thrusting in ltfl. 3. step 3 (fig. 11) illustrates the second phase, during which a hole is formed in the ltfl, and a piggyback imbricate fan of utfl thrust sheets is piled up in the hole. a tentative interpretation is that body forces from the weight of the utfl imbricate fan cause hydrodynamic shear deformation and the extensional normal fault imbricate fan in the ltfl (sf13 in fig. 11). the fault block subject to hydrodynamic fig. 11 four steps in the dynamic development of the glaciotectonic framework sf1 illustrated by balancing and stepwise deformation of the segment s.p. 550–1100 of the seismic section in fig. 3. purple utfl above green ltfl. to calculate the shortening, a fixed pin line is drawn at the front of each thrust-fault level and a loose pin in the rear. step 1 shows the balanced cross-section. the thrust sheets in ltfl are annotated sf11–sf15. step 2 shows initial translation of imbricate thrust faulting in both utfl and ltfl. the thrusting of utfl is slightly ahead of the thrusting in the ltfl. in step 3, the weight of the imbricates in the utfl created a combined hydrodynamic shear deformation and extensional normal fault imbricate with extension about 300 m, see sf13 in the ltfl. in step 4, compressional thrust fault imbrication of the utfl resulted in an imbricate fan. the fan filled the extensional depression and pushed sf12 above the leading thrust sheet sf11 forming a huge hanging wall anticline. the further compression of the ltfl proximal to the depression leads to normal faulting in utfl above sf12. the total compression of this part of the section leads to a shortening of about 3400 m (44%) in the utfl and 600m (9%) in ltfl (see also description in sections 3.3, 4.8 and 4.9). 1000m sf14 sf15 w e step 1 step 2 step 3 step 4 sf11 sf11 sf11 sf12 sf12 sf12 sf13 sf13 sf13 sf14 sf14 sf15 sf15 sf11 sf12 sf13 sf14 sf15 5 x exaggeration shortening 0% 2150m = 28% 400m = 6% 2600m = 34% –300m = 105% 3400m = 44% 600m = 9% 0% hydrodynamic shear front pin purple front pin green rear pin green rear pin green rear pin green rear pin green rear pin purple rear pin purple rear pin purple rear pin purple https://doi.org/10.34194/t407hg84 andersen & pedersen 2025: geus bulletin 59. 8386. https://doi.org/10.34194/t407hg84 14 of 20 geusbulletin.org shear is seen on the seismic sections as a non-reflective area (fig. 3, s.p. 920–850). during this stage, the extension of ltfl was 5% greater than the compression. 4. step 4 (fig. 11) illustrates the final phase, when the excessive weight of the utfl in the hole caused more compression in front, pushing the huge thrust sheet of the ltfl (sf12 in fig. 11). this generated a hanging wall anticline above sf11 and at the same time caused an even greater hole for the piggyback imbricate fan of the utfl to slide into. this movement caused the observed hill-hole pair on the upper décollement surface (fig. 9). hence, it is suggested that body forces from the imbricate fan in utfl were an important factor in the formation of sf1. the resulting compression of the utfl and ltfl is c. 44% and c. 9%, respectively. the structural framework for sf2 is situated in the intermediate domain of ltfl in the northern part of the investigated area (fig. 4, s.p. 0–250; fig. 8). the north-eastern part of the structure could not be fully mapped because the linespacing is too large (figs 1, 8). sf2 consists of an imbricate fan, which, as a whole segment, was translated up along an easterly dipping ramp and displaced across a flat (the intermediate décollement in fig. 4). the top of sf2 is truncated by a roof thrust (the upper décollement). the sequential evolution of sf2 can be divided into four phases, as illustrated in the four steps in fig. 12. these four steps are described as follows: 1. step 1 (fig. 12) illustrates the balanced cross-section of the segment between s.p. 0–440 from fig. 4. the thrust sheets in the ltfl are annotated sf21–sf24. note that the succession with a chaotic reflection pattern above and behind the itfl is excluded from the reconstruction (fig. 4). 2. step 2 (fig. 12) illustrates the early phase of deformation, where the ramp collapses, and imbrication of thrust sheets in the ltfl is initiated. internal deformation of sf24 is happening, and compression of the thrust sheet sf23 results in an imbricate fan. moreover, gentle folding and displacement of thrust sheets in utfl start. 3. step 3 (fig. 12). in this phase, the thrust sheets constituting the imbricate fan (sf23) are pushed farther up the ramp by thrust sheet sf24 and translated piggyback on top of the ltfl towards the foreland. in front of the imbricate fan (sf23), utfl started to form a thrust-fault duplex stack. 4. step 4 (fig. 12). in the final phase, the translation of the imbricate fan (sf23) continues. a combination of uplift and push from the rear resulted in the further development of the duplex stack of the utfl, which is carried piggyback on top of thrust sheets sf21 and sf22. in this final phase, the tailing thrust sheet of utfl became trapped in the position between the translated imbricate fan of the ltfl (sf23) and the duplex stack of utfl. the frontal part of the duplex stack was thrust up and displaced towards the fig. 12 four steps in the dynamic development of the glaciotectonic framework sf2 illustrated by balancing and stepwise deformation of the segment s.p. 0–440 of the seismic section in fig. 4. purple utfl above green ltfl. to calculate the shortening, a common fixed pin line is drawn at the front of the thrust-fault level and a loose pin in the rear of each level. step 1 shows the balanced cross-section. the thrust sheets in ltfl are annotated sf21–sf24. sf23 consists of the imbricate fan itfl (fig. 4). small-scale deformation and hydrodynamically brecciated in the succession above and behind sf23 and sf24 (fig. 4) are excluded from the reconstruction. step 2 shows the initial ramp collapse at the rear end of sf22 and imbrication of ltfl and gentle folding and displacement of utfl above ltfl. in step 3, the imbricate fan (sf23) starts to translate the tailing end of sf22. piggyback on ltfl, the utfl initiated to form a thrust-fault duplex stack above sf22. in step 4, the duplex stack of utfl is pushed by the translated imbricate fan sf23 on the thrust sheet back of sf22, and the tailing thrust sheet of utfl is trapped in the position between the translated imbricate fan sf23 and the duplex stack of utfl displaced towards the foreland. compression is about 2200 m (49%) in the utfl and 3800 m (43%) in the ltfl (see also description in sections 3.3, 4.8 and 4.9). 5 x exaggeration step 1 step 2 step 3 step 4 w e shortening 0% 0% 300m = 7% 1600m = 18% 1600m = 36% 3000m = 34% 2200m = 49% 3800m = 43% rear pin purple rear pin purple rear pin purple rear pin purple rear pin green rear pin green rear pin green rear pin green common front pin 1000 m sf23 sf23 sf22 sf22 sf22 sf24 sf24 sf21 sf21 sf21 sf24 sf23 sf23 sf24 sf22sf21 https://doi.org/10.34194/t407hg84 andersen & pedersen 2025: geus bulletin 59. 8386. https://doi.org/10.34194/t407hg84 15 of 20 geusbulletin.org foreland, and minor ramp thrusting of the frontal part of sf22 accentuated the central depression above the utfl developed in sf2. 5. discussion this study shows that the investigated part of the fanø bugt glaciotectonic complex consists of large-scale thrust fault structures on the scale of hundreds of metres to kilometres. the structures constitute a duplex in two thrust fault levels separated by an internal décollement surface. the basal décollement surface is weakly inclined (around 0.5°) towards the direction of transport of thrust faults. the age of the deformed sediments ranges from lower miocene and younger sediments with a thin quaternary sediment covering above an erosional surface mapped in larsen & andersen (2004). inside the complex, different structural styles are recognised, and in several places, normal faulting occurs, resulting in extension in an otherwise compressional setting. 5.1. unique features of the fanø bugt glaciotectonic complex although glaciotectonic complexes have been described in many studies from all over the world, only a few were discovered on a scale comparable to the fanø bugt glaciotectonic complex (e.g. pedersen & boldreel 2015, 2017; winsemann et  al. 2020; lohrberg et  al. 2022). common amongst those glaciotectonic complexes is that they all involve pre-quaternary sediments. this is possibly because pre-quaternary sediments often by their nature are more compressed than quaternary sediments and, therefore, have greater coherence, hence the ability to evolve larger coherent thrust sheets than seen in studies only involving quaternary sediments (e.g. vaughan-hirsch & phillips 2017; phillips et al. 2018). the heligoland glaciotectonic complex in the german sector is less than 100 km south of fanø bugt glaciotectonic complex and is the most comparable, described by lohrberg et al. (2022). the heligoland glaciotectonic complex is comparable in scale and has many similar structural features, such as imbrications in two levels separated by an internal décollement surface, a detail, which is seldom described for glaciotectonic complexes. analogues in terms of their orogenic setting with multiple detachment levels and imbrications in two levels have been described for the central andes of peru, the norwegian caledonides and the cumberland plateau in the southern appalachians (pfiffner 2017). a basal décollement surface dipping towards the transport direction of the thrust faults and the ice sheet, as seen in the fanø bugt glaciotectonic complex, is also found at the heligoland glaciotectonic complex (lohrberg et al. 2022) and in other glaciotectonic complexes in the eastern part of the north sea (huuse & lykke-andersen 2000a). the basal décollement of the fanø bugt glaciotectonic complex is situated in lower miocene sediments and has a maximum depth of 400 ms twt. the basal décollement of the heligoland complex is situated a bit higher (350 ms twt) and at a younger stratigraphic level, the miocene-pliocene boundary (lohrberg et  al. 2022). the glaciotectonic complexes situated north of fanø bugt have décollements at the mid miocene unconformity or at the base quaternary boundary at a depth of 150–300 ms twt (huuse & lykke-andersen 2000a). three zones of deformation are identified in both the heligoland glaciotectonic complex (lohrberg et al. 2022) and the fanø bugt glaciotectonic complex. however, while the heligoland glaciotectonic complex shows the steepest thrust angles close to the ice margin in the hinterland, like ones observed in other glaciotectonic complexes (e.g. pedersen 2005; phillips et al. 2017), the opposite pattern is observed in the ltfl in the fanø bugt glaciotectonic complex. here, thrust faults steepen towards the foreland (fig. 3, no exaggeration section), deviating from the typical pattern observed in fold and thrust deformation. the average shortening found in this study varies from 9 to 49% (figs 11 and 12). it was measured not only across parts of the complex that potentially experienced the largest compression but also in the area of largest extension (ltfl in sf1, step 3; figs 3–6, 8). hence, it only resembles the shortening in the part of the complex where the two profiles exist and does not represent the total shortening of the fanø bugt glaciotectonic complex. this study shows that the ltfl has experienced much less compression in certain areas than the utfl and also less compression when compared to glaciotectonic complexes in general, 11–61% (values from croot 1988; hart 1990, 1994, 1995; klint & pedersen 1995; pedersen 1996, 2005, 2014; harris et al. 1997; boulton et al. 1999; huuse & lykke-andersen 2000a; williams et  al. 2001; gehrmann 2019; winsemann et al. 2020; lohberg et al. 2022; vaughan et al. 2024). this is most likely due to the relatively large amount of extension found in the ltfl in certain areas of the fanø bugt glaciotectonic complex. furthermore, the limitations of the seismic data in resolving small-scale or very steeply inclined structures will result in a minimum calculation of the compression, when compared to studies on seismic data with a higher resolution or studies of onshore outcrops where smallscale details can be included in the restoration of the glaciotectonic complexes. we attribute the unique features of the fanø bugt thrust complex (highlighted here) to the deformation https://doi.org/10.34194/t407hg84 andersen & pedersen 2025: geus bulletin 59. 8386. https://doi.org/10.34194/t407hg84 16 of 20 geusbulletin.org mechanism. thus, there may have been a subtle interaction between various features that could facilitate deformation during the formation of the complex. 5.2. deformation mechanisms and structural styles in the fanø bugt glaciotectonic complex, different deformation domains exist, each characterised by various types of structures, a common phenomenon in thinskinned tectonics both in orogenic settings and glacial settings (e.g. boulton et  al. 1999; pedersen & boldreel 2015). various structural styles occur in the individual thrust fault levels, which may be related to the former position of the ice-sheet margin responsible for the deformation. glaciotectonic deformation is generally considered to be driven by a gravity-spreading mechanism (e.g. van der wateren 1985; pedersen 1987; aber et al. 1989; andersen et al. 2005); however, it could also be argued that some of the structures found in this study indicate a gravity-gliding mechanism contributed to the deformation. the utfl can be divided into two different structural domains based on the different structural styles: (1) a distal domain showing thrusting and folding implying proglacial deformation; (2) a proximal domain showing a chaotic reflection pattern, interpreted to resemble pervasively sheared or chaotically mixed sediments (aber 1985), boudinage structures (berthelsen 1979), fluid flow structures and hydrodynamic brecciation (pedersen 2005). these structures imply subglacial deformation or a combination of both proglacial and a later phase of subglacial deformation. the reflection pattern of the strongly sheared and deformed sediments of the utfl stops abruptly at the upper décollement surface (fig. 3, utfl, proximal domain). hence, it is suggested that the upper décollement surface absorbs the induced subglacial stress and prevents the same kind of deformation of the sediments below this surface. an exception is seen in the easternmost part of the study area, where the entire sediment package down to the basal décollement surface was probably subject to strong subglacial deformation (fig. 3, ltfl, s.p. 0–125). the ltfl exhibits three structural domains with different structural styles, which changes in dip (fig. 3) and strike (fig. 8) of the faults along the direction of transport. the structural styles mostly imply proglacial deformation, except for the presence of extensional faulting, mainly in the intermediate domain (fig. 3), and the pervasively sheared or chaotically mixed sediments in the easternmost end of the proximal domain (fig. 3). the structural frameworks in the intermediate domain comprise a combination of compressional and extensional faulting, which may imply subglacial deformation close to the ice margin (croot 1987). hence, each framework may reflect the position of the ice margin. if this is the case, subglacial stress did go deeper than the upper décollement. however, the excessive load of the utfl piggyback imbricates in the hole of sf1 could also have induced the hydrodynamic shear and extension, as seen in the ltfl (sf13, step 3, fig. 11). extensional faulting in glaciotectonic settings has been described in a few studies (croot 1988; pedersen 2005; vaughan et al. 2024). however, the existence of several structural frameworks of thrust faults with tailing listric normal faults detaching on a gently dipping basal décollement surface (fig. 3, intermediate domain of ltfl) has, to our knowledge, not previously been reported in glaciotectonic settings. this structural framework has, although on a much smaller scale, huge similarities with structural frameworks seen in deepwater thrust fault belts (dwtfb; e.g. morley et al. 2011; mahanjane & franke 2014), where a linkage of up-dip extension with down-dip contraction via a dipping detachment zone is usually associated with the mechanism of gravity gliding (morley et al. 2011). this combination of structural styles and deformation types seems to be unique for the fanø bugt glaciotectonic complex, compared with other large-scale glaciotectonic complexes. one hypothesis is that the deformation mechanism was gravity spreading caused by the load of an ice sheet, and in the case of sf1 helped by the excessive body mass of a piggyback imbricate fan in the utfl, combined with the mechanism of gravity gliding due to the inclined basal décollement surface. the basal décollement surface in the fanø bugt glaciotectonic complex is a gently inclined surface situated in lower miocene sediments with an approximate dip of 0.5° (andersen 2004). a distinct change in strike and dip of the basal décollement surface occurs on each side of the study area (fig. 7). hence, the area stands out as having a lower inclination than the surrounding areas and a dip in the main part of the area parallel to the direction of transport of the ltfl. this begs the question as to whether this variation in dip of the décollement surface could have influenced the deformation. one hypothesis is that porewater flowing from the north was directed towards the study area following the dip of the basal décollement surface. the decrease in dip and change in strike on the border of the study area (fig. 7) would slow porewater flow, allowing porewater pressure to build up. the high porewater pressure could have facilitated deformation and enhanced the possibility of both gravity spreading and gravity gliding. the fanø bugt glaciotectonic complex was formed by an ice sheet coming from an eastern source area (huuse & lykke-andersen 2000a; andersen 2004). the https://doi.org/10.34194/t407hg84 andersen & pedersen 2025: geus bulletin 59. 8386. https://doi.org/10.34194/t407hg84 17 of 20 geusbulletin.org coincidence in the dip direction of the basal décollement surface and transport direction of the ice sheet possibly enhanced the deformation forces and caused the deformation to reach its maximum extent in the area studied here compared to the rest of the fanø bugt glaciotectonic complex, note the curvature on the deformation front (fig. 7). the dipping décollement towards the foreland in the fanø bugt complex contrasts with the common critical taper wedge models, where the basal detachment dips in the opposite direction towards the hinterland (davis et al. 1983). however, it is important to keep in mind that it is the direction of dip of the topographic surface that drives the thrust movement and not the dip of the décollement (van der pluijm & marshak 2004). in the case of a glaciotectonic complex, the topographic surface of the ice provides the driving mechanism (aber et  al. 1989; andersen et al. 2005). in the case of the fanø bugt glaciotectonic complex, a relatively huge topographic high must have developed in connection with the formation of sf2, where the green reflection is elevated about 150 ms (c. 140 m) above normal level (fig. 4). however, this raised topography is concentrated in a relatively small area (fig. 8). 5.3. the kinematics of the fanø bugt glaciotectonic complex the kinematic evolution of the two structural frameworks sf1 and sf2 shows that the deformation of the utfl preceded that of the ltfl (figs 11 and 12). this is supported by the map of the internal décollement (fig. 9), which shows a wavy pattern with highs above thrust ridges and lows in between, indicating superimposed deformation. this should be seen in contrast to the area south of the study area, where the utfl is sliding on a planar décollement surface, and no ltfl exists (figs 5 and 6). based on our observations, we suggest that progressive deformation took place nearly simultaneously, in the ltfl and utfl, but with a small difference in timing during propagation from the east to west. the deformations reach their maximum depth in the south-western corner of the study area, approximately coinciding with the 400 ms twt (c. 360 m) contour curve of the basal décollement (fig. 7). this may indicate that the south-western limit of the deformation of the ltfl in the fanø bugt glaciotectonic complex was controlled by the depth to the basal décollement (andersen 2004; andersen et al. 2005). the deformation of the utfl continues farther to the south beyond the study area (figs 5 and 6). the westward progression of the deformation front is located at the same position for both the utfl and the ltfl (figs 3, 4, 10), suggesting that the deformation at both levels occurred during the same deformation event. this conclusion was also drawn in the case of the heligoland glaciotectonic complex (lohrman et  al. 2022), where similar thrust fault architecture is observed. a key argument, which applies to both the heligoland and fanø bugt glaciotectonic complexes, is the absence of undisturbed sediments and erosional truncation between the two levels of deformation. this suggests that the folds and thrusts at both levels formed contemporaneously, likely as part of a single deformational event. 6. conclusions the mapped area of the fanø bugt glaciotectonic complex contributes significant documentation of largescale glaciodynamic features in the south-eastern north sea. the complex architecture includes two main thrustfault levels (ltfl and utfl) and possibly a third (itfl) in a small area. the ltfl is a duplex comprising imbricate thrust sheets and normal faults, detaching on a westward dipping basal décollement surface. the roof thrust of the duplex acts as an internal décollement, separating imbricated thrust faults of the utfl from the ltfl. thrust ridges range from hundreds of metres to kilometres scale, with crest elevation varying from a few tens of metres to more than 150 m above the level of the foreland deposits. deformation extends down to c. 400 ms twt at c. 360 m b.s.l., and prominent hanging-wall ridges can be traced along strike for more than c. 10 km. the utfl reflects sequential superimposed deformation of the ltfl, both presumably formed during the same pre-weichselian ice advance from the east. the deformation propagated westward with slight temporal offsets between the ltfl and the utfl and a varying shortening within the complex. structural styles vary across the complex, from east to west, as do the strike and dip of the faults. the utfl generally deforms in a more ductile way than the ltfl, due to differences in sediment rheology. the ltfl represents the early/middle miocene marine arnum formation with possible intercalations of the sandy bastrup and odderup formations, whereas the utfl constitutes the middle miocene, fully marine clayey hodde and gram formations. two spectacular structural frameworks, sf1 and sf2, have been analysed. both structures include a hidden hill-hole pair potentially separated by a tear fault. a tentative restoration of profiles across sf1 and sf2 shows an average minimum shortening of c. 9–43% of the ltfl and c. 44–49% of the utfl. the displacement varied in the order of –300 m (extension) and up to 3.8 km (compression). the unique combination of architecture and structural styles suggests a deformation mechanism involving https://doi.org/10.34194/t407hg84 andersen & pedersen 2025: geus bulletin 59. 8386. https://doi.org/10.34194/t407hg84 18 of 20 geusbulletin.org a combination of proglacial gravity spreading caused by the differential load of the ice sheet, possibly aided by gravity gliding due to the inclined basal décollement surface. high porewater pressure and the alignment of ice-sheet movement with the basal décollement dip possibly facilitated deformation further. if an opportunity arises to collect core material from within the complex, it would be valuable for future research to enable detailed analyses of the physical and mechanical properties of the sediments in and around the décollement surfaces and thereby enhance our understanding of the mechanism behind their formation. core material could also help constrain the glaciostratigraphic timing of the complex. furthermore, seismic data with a higher resolution could contribute with images of smaller-scale structures, distinct sedimentary units and more accurate erosional contacts within the top 50–100 ms twt, hence provide better conditions for reconstructing a detailed history of ice advance. acknowledgements holger lykke-andersen, peter r. jakobsen and james a. chalmers are thanked for valuable discussions and comments on a first draft of the manuscript. kristian a. rasmussen and anders mathisen are thanked for the help with the topographic subsurface maps and jette halskov for the help with refinement of figures. the authors would like to thank david collin tanner and an anonymous reviewer for providing critical comments and suggestions that greatly improved the manuscript. additional information funding statement this project was financed by the geological survey of denmark and greenland and the danish research agency. the seismic data were financed by the royal danish administration of navigation and hydrography, the geological survey of denmark and greenland and the university of aarhus. the royal danish navy provided the ship for acquisition. author contributions lta: conceptualisation, interpretation of seismic data, writing original draft, structural analysis, visualisation. sasp: conceptualisation, interpretation, structural analysis. competing interests the authors 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(ed.): investigation in geophysics, volume 2. society of exploration geophysicists investigations in geophysics 2, 526 pp. tulsa, oklahoma. https://doi.org/10.34194/t407hg84 https://doi.org/10.34194/geusb.v33.4479 https://doi.org/10.1002/jqs.2887 https://doi.org/10.1002/jqs.2887 https://doi.org/10.3390/geosciences7030071 https://doi.org/10.3390/geosciences7030071 https://doi.org/10.1016/j.quascirev.2017.03.027 https://doi.org/10.1016/j.quascirev.2017.11.001 https://doi.org/10.1016/j.quascirev.2019.105943 https://doi.org/10.1016/j.quascirev.2019.105943 https://doi.org/10.37570/bgsd-1996-43-14 https://doi.org/10.37570/bgsd-1996-43-14 https://doi.org/10.37570/bgsd-2004-51-07 https://doi.org/10.37570/bgsd-2004-51-07 https://doi.org/10.34194/geusb.v22.4733 https://doi.org/10.34194/geusb.v22.4733 https://doi.org/10.34194/seriec.v12.7106 https://doi.org/10.34194/seriec.v12.7106 https://doi.org/10.34194/geusb.v38.4388 https://doi.org/10.13140/rg.2.1.4118.2803 https://doi.org/10.13140/rg.2.1.4118.2803 https://doi.org/10.1016/j.quascirev.2013.03.016 https://doi.org/10.37570/bgsd-1985-34-06 https://doi.org/10.37570/bgsd-1985-34-06 https://doi.org/10.1016/j.pgeola.2024.04.001 https://doi.org/10.1016/j.pgeola.2024.04.001 https://doi.org/10.1002/jqs.2836 https://doi.org/10.1111/bor.12461 https://doi.org/10.1144/jgs.158.1.125 https://doi.org/10.1016/j.quascirev.2019.106068 https://doi.org/10.1016/j.quascirev.2019.106068 piggyback imbrications, duplex stacking and sequential superimposed deformation in the fanø bugt gla 1. introduction 2. geological setting 3. data and methods 3.1. seismic data acquisition and parameters 3.2. interpretation, maps and plots 3.3. structural analysis 4. results 4.1. the seismic interpretation 4.2. the mapped surfaces 4.3. architecture and structural styles 4.4. the basal décollement surface 4.5. the structural trend of the lower and intermediate thrust fault levels 4.6. the upper décollement surface 4.7. the structural trends in the utfl 4.8. restoration of the deformed sections 4.9. the structural framework and dynamic development of sf1 and sf2 5. discussion 5.1. unique features of the fanø bugt glaciotectonic complex 5.2. deformation mechanisms and structural styles 5.3. the kinematics of the fanø bugt glaciotectonic complex 6. conclusions acknowledgements additional information funding statement author contributions competing interests data availability statement references figures fig. 1 location in the south-east danish north sea of the fanø bugt glaciotectonic complex and the fig. 2 stratigraphic diagram (modified from rasmussen 2004). the succession involved in the thrustin fig. 3 seismic section (fl01-50) showing a duplex with imbricated thrust sheet in two levels. the lo fig. 4 seismic section (fl01-54) showing an approximate dip profile across the structural framework fig. 5 seismic section (fl00-27) is crossing the area oblique to strike of the structural elements a fig. 6 seismic section (fl01-19) is crossing the area oblique to the strike of structural elements a fig. 7 a two-way time-contour map of the basal décollement surface and its continuation beyond the w fig. 8 a structural two-way time (twt) contour map showing the strike of the main structures at t fig. 9 3d map of the upper décollement surface with a 2.5 times exaggeration gives a perspective vie fig. 10 map comparing the trends of the faults of the utfl (black) compared with those of the ltfl ( fig. 11 four steps in the dynamic development of the glaciotectonic framework sf1 illustrated by bal fig. 12 four steps in the dynamic development of the glaciotectonic framework sf2 illustrated by ba voutchkova et al. 2024: geus bulletin 57. 8374. https://doi.org/10.34194/geusb.v57.8374 1 of 10 data article | short open access nationwide data sets for drinking water hardness at public waterworks and their water supply areas in denmark denitza d. voutchkova*1 , charlotte t. thomsen2 , niels claes3 , lars a. olsen2 , lærke thorling1 , bjarni pjetursson2 , birgitte hansen1 1department of geochemistry, geological survey of denmark and greenland (geus), aarhus, denmark; 2geological data centre, geological survey of denmark and greenland (geus), copenhagen, denmark; 3aarhus vand, viby, denmark three spatiotemporal data sets of drinking water hardness in denmark (version 1) are presented here: (1) annual drinking water hardness at public waterworks (1905–2023); (2) annual drinking water hardness at their water supply areas (1978–2023) and (3) the latest drinking water hardness at the water supply areas (1980–2023). raw data were extracted from the jupiter database for groundwater and drinking water data in denmark, and were quality-assured. hardness was calculated after semi-automatic outlier exclusion based on ca and mg, or if not available, the reported total hardness. data were further aggregated at the waterworks level by the annual mean and at the supply area level by the weighted mean (weighted to waterworks annual abstraction volumes). temporal and spatial gaps were filled prior to these aggregations. various stakeholders could benefit from these open access data. they provide a societal service in response to increased public interest in drinking water hardness. the research community could use the data in environmental, exposure or epidemiological assessments. finally, the water supplies and the public sector could benefit from these data as they provide a nationwide overview of current and past drinking water hardness in denmark and highlight the geographic areas that lack recent data, most probably due to de-regulation. *correspondence: dv@geus.dk received: 05 mar 2024 revised: 07 may 2024 accepted: 12 may 2024 published: 05 sept 2024 keywords: drinking water hardness, calcium, magnesium, public waterworks, water supply areas abbreviations: dw: drinking water geus: geological survey of denmark and greenland lod: limit of detection sd: standard deviation wsa: water supply area geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: adam hambly (dtu, denmark) reviewed by: mayumi hori (university of tokyo, japan), ronnie levin (harvard university, usa), camilla tang (niras, dk) funding: see page 9 competing interests: see page 9 additional files: see page 9 tabular abstract geographical coverage denmark temporal coverage 1905–2023 subject(s) engineering and environmental geology, geochemistry, cosmochemistry and geochronology, geoscience history and policy data format(s) analysed data (shape files and csv files). sample collection & analysis public waterworks metadata and drinking water analyses of ca, mg and total hardness from the jupiter database extracted on 6 september 2023. the water supply areas are from schullehner (2022,  https://doi.org/10.34194/geusb.v49.8319). data were quality assured, filtered, aggregated and linked together. parameters drinking water hardness in german degrees (°dh), drinking water hardness class, unique identification number for the water supply areas (wsaid) and for the public waterworks (anlaegid), as well as sampling year, based on the sampling date (proevedato). data availability all data files are available at https://doi.org/10.22008/fk2/nlx5sx potential application(s) for these data data set 1 (annual hardness at the waterworks level for the period 1905–2023) allows for in-depth spatiotemporal analyses at different scales – from time-series analysis at single waterworks to national status overviews and spatial pattern analysis. data set 2 (annual hardness data at the water supply areas, wsas for the period 1978–2023) allows for exposure and epidemiological studies on a national scale by spatial link with geocoded household addresses or patients residential history. data set 3 (latest hardness at wsa level) can be used for mapping the current drinking water hardness in denmark and for visualising current data availability. these three data sets have many other potential applications spanning research, education, outreach and public decision support. https://doi.org/10.34194/geusb.v57.8374 https://orcid.org/0000-0003-2840-072x https://orcid.org/0000-0002-0030-2418 https://orcid.org/0000-0003-0713-0214 https://orcid.org/0000-0002-7873-9572 https://orcid.org/0000-0001-9067-0005 https://orcid.org/0000-0002-0685-0941 https://orcid.org/0000-0003-2318-145x mailto:dv@geus.dk https://creativecommons.org/licenses/by/4.0/deed.ast https://doi.org/10.34194/geusb.v49.8319 https://doi.org/10.22008/fk2/nlx5sx voutchkova et al. 2024: geus bulletin 57. 8374. https://doi.org/10.34194/geusb.v57.8374 2 of 10 geusbulletin.org data collection background information the main purpose of this work is to update the national drinking water (dw) hardness map for denmark produced by the geological survey of denmark and greenland (geus 2010). the previous map was based on data from 2010, now outdated, and was aggregated and presented at the municipality level. this approach is problematic for at least two reasons. firstly, some municipalities are supplied by waterworks located outside of the municipality borders (e.g. in the copenhagen area and odense). this results in both gaps in the map for these municipalities and miscalculations for the municipalities where the waterworks are located (but do not supply water locally). secondly, some of the large municipalities in jylland and on sjælland have dw with three different hardness-classes, so if municipality-level averaging is used, the resulting dw hardness would be significantly different from that experienced by consumers. interpolation based on waterworks location is equally problematic because it assumes that the dw is supplied locally in the surroundings of the waterworks, which is not necessarily the case. therefore, we decided that the optimal way to represent dw hardness is currently at the water supply area (wsa) level, using the open access data set by schullehner (2022). the wsa data were collected at the municipality level, digitised and linked to the waterworks they are supplied by schullehner (2022). the data cover the period 1978–2019 (schullehner 2022). dw quality data, including data on calcium (ca), magnesium (mg) and total hardness, is collected to comply with dw quality regulations (council of the european union 2020; miljøministeriet 2022, 2023a). all compliance samples are taken, analysed and reported to the national well database jupiter (hansen & pjetursson 2011) by accredited chemical laboratories in denmark. the legal requirements for the quality of environmental samples (miljøministeriet 2023b) must be followed, according to which, the relative uncertainty for ca and mg must be ≤ 15% (for  high concentrations) and the absolute uncertainty (for low concentrations) must be ≤ 3 or 1 mg/l, respectively. these requirements have been in force since their introduction in 2002. method m069 of the danish environmental protection agency reference laboratory for chemical and microbiological measurements provides further information on the appropriate pre-treatment, storage and laboratory methods (miljøstyrelsen 2023). the legal requirements for compliance sampling for ca, mg and total hardness have changed over the years and this is reflected in data availability in jupiter. since the end of 2017 (miljøministeriet 2017), there is only a requirement to provide a general description of the dw quality to the consumer, including the water hardness (§ 33(6), miljøministeriet 2023a) but there are no longer dw quality standards for ca, mg and total hardness, and thus no required sampling frequency based on the waterworks’ capacity (m3 of produced dw). this is in fact a de-regulation, as for the period 2001–2017, ca, mg, and total hardness were included in the extended danish dw control list (miljøministeriet 2001, appendix 5). hence, they were sampled once every year or every second year, depending on the waterworks’ capacity. in this period, there was a quality standard for mg (50 mg/l), a remark that ca should not exceed 200 mg/l and that the total hardness should be 5–30 german degrees (°dh; miljøministeriet 2001, appendix 1a). before that (1988–2001) ca, mg, and total hardness were also in the extended dw control (miljøministeriet 1988, appendix 5), but the sampling frequency was once a year irrespective of waterworks’ capacity. there was no explicit standard for ca, only a note referring to the standard for total hardness (5–30 °dh) and next to the standard for mg (50 mg/l) there was also a provisional guideline value (30 mg/l; miljøministeriet 1988, appendix 1). there is considerable public interest in the dw hardness, as it is an important technical characteristic of tap water in relation to precipitation of carbonates (i.e. scaling) in water pipes, boilers and other hot water household appliances, as well as in relation to increased soap use. the data have also potential relevance to the public health sectors and various industry sectors. data sources raw data were extracted from the jupiter database (hansen & pjetursson 2011) on 6 september 2023. field (column) names and codes are capitalised, italicised and provided in parentheses, where relevant. a list of field names used in these data sets and their corresponding field name in the jupiter database with their pc jupiter xl naming format is provided in table s1 (supplementary file s1). figure 1 summarises the processing steps used in creating the data sets. database tables containing information on the dw samples, their chemical analyses and waterworks were all joined together by the unique sample identification number (id) and the waterworks id (fields proeveid and anlaegid, respectively). only waterworks with chemical analyses were retained in the data set. annual groundwater abstraction volumes (estimmaengde) were joined to the data set using the waterworks id and the sampling year (fields anlaegid and proeveaar, respectively). data-quality checks and filtering procedures described here were executed before combining the input tables (provided in supplementary files s3). https://doi.org/10.34194/geusb.v57.8374 http://www.geusbulletin.org/ voutchkova et al. 2024: geus bulletin 57. 8374. https://doi.org/10.34194/geusb.v57.8374 3 of 10 geusbulletin.org data preparation quality checks and filtering procedures used for preparing the data sets are similar to those used by voutchkova et al. (2021). only public waterworks were included. public waterworks supply more than 10 households and are registered in jupiter as ‘company type’ (virksomhedstype field) as follows: • municipality-owned public water supplies (v01) • privately owned public water supplies (v02) • water supply (m42) – a code used in the past for public water supplies. the most recently registered value for ‘company type’ (virksomhedstype) was used. only the analyses for ca, mg and total hardness were retained in the data set (stofnr = 280, 321, 70, where stofnr is the field name referring to the chemical parameter, according to stancode). the reported total hardness was used to fill gaps where ca or mg were unavailable for the hardness calculation and for comparison purposes. hardness was calculated based on ca and mg, due to their overall higher spatiotemporal data coverage than total hardness (see section on statistical analysis and data processing). the units for ca and mg (field name enhed = 1, i.e. mg/l) and total hardness (enhed = 22, german degrees, °dh) were checked. data without a reported unit were excluded. flagged data were excluded based on the attribut field unless the field was empty or contained ‘<’ (i.e. below the limit of detection, lod). samples not approved by the owner were indicated in the kvalitetssikring field and excluded from the data sets. all samples from treatment processes (process waters, indicated by projekt = ‘procv’) were excluded, as they are not representative of the finished treated dw. depending on the sampling purpose, only dw samples were kept in the data set where they had the following formaal values: 0, 6, 7, 9, 10, 11, 12, 34, 35, 39, 40, 41, 42 (definitions for these codes are available in the documentation of pc jupiter xl; geus 2024). a formaal of nine refers to samples taken for the purpose of pollution assessments, however since the sampling location was either at the exit waterworks or at the tap, they are none-the-less dw samples. so, 79 samples with this code were retained. samples with no value entered for formaal were excluded (n = 319). as a result, only samples taken at exit waterworks, the supply network and at the end-consumers’ tap affiliated with a specific waterworks were retained in the data set. ca and mg reported as 0 were treated as reporting errors and excluded. concentrations in field maengde must be reported to jupiter as > 0. for analyses below the lod, the lod must be provided in the maengde field, and the attribut field must contain ‘<’. the other jupiter lod field often contains incomplete and contradictory information, so it was not used here. a check for unreasonably high lods was also performed. for ca, two samples with lod = 97.8 mg/l and 132 mg/l were most probably a result of erroneously used attribut ‘<’ and were excluded from the data set. subsequently, lods for ca ranged from 0.005 to 5 mg/l and only 10 samples were below the lod. for mg, lod ranged from 0.005 to 9.4 mg/l (n = 44) with the most frequent lod = 1 mg/l (n = 18). no samples were excluded for mg based on high lod. the attribut ‘<’ was further ignored, i.e. < lod = lod. data with erroneous sampling dates (field name proevedato; e.g. 0001, 0208, 1894 or 1897) were excluded. only samples in the period 1900–2023 were kept in the data set. jupiter contains spatial information about the location of the waterworks. projected coordinates (utm 32, euref 89) were used here (field names xutm32euref89 and yutm32euref89). in 2019, schullehner (2022) fig. 1 flowchart summarising the data preparation and statistical analyses used here. qa/qc: quality assurance and quality control. wsa: water supply area. for the contents of data sets 1, 2, and 3, see section ‘data description and main features’. hardness calculated by eq. (3). field names: anlaegid: unique identification number (id) of the waterworks. plantid: waterworks id used in schullehner (2022), same as anlaegid. proveid: sample id. proeveaar: sample year. wsaid: id of the wsa (schullehner 2022). full list of field names in supplementary files s1, s2. jupiter 3963 waterworks 1905–2023 qa/qc & outlier removal plantid database processing step input/output symbols hardness eq. (3) sample-mean annual-mean at waterworks data set 1 proeveid proeveaar & anlaegid extraction 6 sep. 2023 wsa join anlaegid mean at wsa proeveaar & wsaid 3571 wsa 1978–2023data set 2 filter schullehner (2022) max(proeveaar) for each wsaid data set 3 3827 waterworks 3571 wsa 1978–2023 https://doi.org/10.34194/geusb.v57.8374 http://www.geusbulletin.org/ voutchkova et al. 2024: geus bulletin 57. 8374. https://doi.org/10.34194/geusb.v57.8374 4 of 10 geusbulletin.org checked, corrected and geocoded missing coordinates of public waterworks where an address was available. these adjusted coordinates were used here (also in supplementary files s3). zero coordinates were substituted with missing values. there were 64 waterworks with missing x or y coordinates (17 861 individual ca and mg analyses). an address was provided for 24 of these waterworks, but geocoding was not done. for waterworks without x or y coordinates, but with an overlying waterworks (i.e. where the overanlaeg field was not empty), the waterworks id was overwritten with the id of the overlying waterworks. in this way, the reported chemical analyses for the waterworks without coordinates were transferred to the overlaying waterworks with coordinates. in addition, seven waterworks (anlaegids 106730, 193235, 189193, 190517, 191992, 192278, 192302) had erroneous coordinates outside denmark, determined after mapping in qgis. for two of those (anlaegids 192278 and 192302) there was a value in the overanlaeg field, so their analyses were transferred to the specific overanlaeg field. we recommend addressing issues with missing or wrong coordinates in future updates of these data sets. identical replicates were removed from the data and outliers in the ca, mg and total hardness time series were excluded by the semi-automated outlier detection procedure. statistical analysis and data processing semi-automated outlier detection procedure the detection and removal of outliers was deemed necessary based on cumulative distribution plots (see fig. s1, supplementary file s1). a semi-automated outlier detection procedure was implemented for each parameter separately (ca, mg and reported total hardness). the outliers were evaluated based on the time series for the specific waterworks. laboratory analyses were flagged as outliers where they were different from their surrounding values. potentially high outliers had a percentage difference ≥200% as calculated in eq. (1), from the two previous and two following analyses: ± = − ×± ± di j x x x 100%i j i i j (1) where, xi is the chemical analysis in the timeseries of a specific waterworks that is evaluated for being an outlier, i is the time index for that chemical analysis, and j is equal to 1 or 2, as it indicates the position of the preceding or following chemical analyses. for example, when j = 1, xi−1 is the chemical analysis before xi , and when j  =  2, xi−2 is the laboratory analysis before xi−1 therefore, the percent difference from the previous one, two, or next one, two analyses were calculated (di±j). values were flagged as potential outliers as follows: for xi , if all di±j ≥ 200% →xi = potential outlier (2) when xi   appeared at the beginning or end of the time series, the corresponding di±j could not be calculated, so only the rest of the conditions were used. for example, for the first analysis in the time series, the difference was calculated based solely on the following two samples. it was assumed that a change in treatment (e.g. softening) would result in a step change followed by a relatively stable period, and thus not all four conditions in eq. (2) would be satisfied simultaneously. however, there could be other reasons for false detection of outliers, for example if the waterworks blends groundwaters or treated dw with very different quality and the mix is not consistent in time. therefore, to validate the outlier detection all time series with flagged potential outliers were plotted and visually examined. while visually examining the time series, outliers in the low concentration or hardness range were discovered. equation (1) flagged the values surrounding them (but not all four conditions from eq. (2) were simultaneously fulfilled). so, to pinpoint the low outliers, the flags were shifted back or forward once, or twice, depending on the specific condition. outlier detection at the end of the time series (the most recent one or two analyses) was not implemented as there was a higher chance that outliers here were due to (1) changes in treatment, (2) introducing new wells with different water quality or (3) due to other operational changes, which could not be evaluated currently. therefore, there were at least two analyses following a potential low outlier. all validated outliers (both high and low) were excluded from the data set, see ‘outlier detection and validation’ section in supplementary file s1 for details. future work could focus on developing a fully automated outlier detection and removal method and testing the performance of different thresholds. dw hardness calculation and classification after outlier removal, the data were considered clean. a sample mean for each parameter was calculated (using fields proeveid and stofnr) as sampling replicates with different concentrations were not treated as individual samples. dw hardness was calculated at the sample level by eq. (3) following hansen & thorling (2018): ° = × +     + + dh ca mg5.6 40.1 24.3 2 2 (3) where  °dh is the dw hardness in german degrees, and ca2+ and mg2+ are the concentrations of ca and https://doi.org/10.34194/geusb.v57.8374 http://www.geusbulletin.org/ voutchkova et al. 2024: geus bulletin 57. 8374. https://doi.org/10.34194/geusb.v57.8374 5 of 10 geusbulletin.org mg in mg/l. the °dh is defined as 10 mg/l cao. it can be converted to american degrees of hardness (i.e. ppm or 1 mg/l caco3) with a conversion factor of 17.848. it was decided to use eq. (3), because overall, ca and mg have a better spatiotemporal coverage over the reported total hardness. figure s6 (in supplementary file s1) shows a comparison between reported and calculated hardness. the dw hardness was calculated based on eq. (3) for 91 673 samples (3939 waterworks). in addition, 2980 samples reported total hardness but no ca and/or mg, and eq. (3) could not be used (fig. s7, supplementary file s1). these were added to the data set to ensure maximum data coverage. thus, dw hardness was either calculated or reported directly for 94 653 samples at 3963 waterworks in the period 6 february 1905–28 august 2023. dw hardness in denmark is traditionally classified in seven hardness classes from very soft to extremely hard (table 1). various classifications are used internationally (rygaard & albrechtsen 2020), so here we report both the hardness classes and the hardness in °dh. we recommend using the colour palette in table 1, because it is (1) suitable for people with deuteranomaly, protanopia, deuteranopia and tritanopia (tested with the tool developed by nichols n.d.) and (2) neutral, without implicit negative associations for danger or bad quality. data aggregation at the waterworks level annual mean dw hardness was calculated for each waterworks and rounded to one decimal (based on the proeveaar and anlaegid fields). this aggregation was necessary because the sample frequency varied from once each 4 or 5 years to multiple times a year. to some extent, this heterogeneity is caused by the legal requirements for sampling frequency (see section ‘background information’). data set 1 contains the resulting annual dw hardness (n = 76 570) for 3963 public waterworks between 1905 and 2023. link between waterworks and wsas data set 1 was joined to the wsa data set (schullehner 2022) with the waterworks id (‘anlaegid’ in our data set or ‘plantid’ in schullehner 2022). some wsas are supplied by multiple waterworks, thus the join type was one-to-many. it was also specified that the sampling year (proeveaar) should be within the period of validity of the wsa (proeveaar = [start, end]). the wsas were last updated in 2019 (schullehner 2022), so here we assumed that they have remained unchanged (i.e. for proeveaar > max(end), the wsas with corresponding end = 2019 was used). this assumption may not hold after 2019 where there has been (1) introduction of new waterworks or (2) consolidation of waterworks (closing of small waterworks) or (3) expansion of wsas due to newly built areas, previously outside wsas. joining the data sets resulted in 3827 waterworks (96.6% of those from data set 1) being linked to 3571 wsas. we recommend that the data set by schullehner (2022) should continue to be updated for use in future assessments. estimating the dw hardness at wsas the annual dw hardness at the wsa level (°dhwsa) was calculated as a volume-weighted mean using eq. (4): dh v dh vwsa° = ∑ × ° ∑     n n n (4) where n is index for the waterworks within a specific wsa, vn is the groundwater abstraction volume for the specific waterworks and year, and °dhn is the annual dw hardness for the specific year from data set 1. when n = 1, the annual dw hardness at the wsa was equal to the annual dw hardness at that waterworks. in total, 660 wsas were supplied by more than one waterworks. the volume-weighted mean was used to account for the different capacities of waterworks, because there could be a few orders of magnitude difference in the estimated annual abstraction. this approach was also used in geus (2010), where the volume-weighted mean was calculated on a municipality level. the abstraction volume data were incomplete (see section ‘estimating annual production volumes’), so calculating a volume-weighted mean introduces uncertainty in the annual dw hardness estimation. for purposes of comparison, an arithmetic mean and standard deviation (sd) were calculated for each wsa (where n > 1) and included in data set 2 together table 1 drinking water (dw) hardness classification and recommended colour palette to represent each class. hardness (°dh) hardness class colour palette code original danish english translation hex rgb ≤ 4 meget blødt very soft #ffffcc 255, 255, 204 4–8 blødt soft #c7e9b4 199, 233, 180 8–12 middelhårdt medium hard #7fcdbb 127, 205, 187 12–18 temmeligt hårdt fairly hard #41b6c4 65, 182, 196 18–24 hårdt hard #1d91c0 29, 145, 192 24–30 meget hårdt very hard #225ea8 34, 94, 168 > 30 særdeles hård extremely hard #0c2c84 12, 44, 132 https://doi.org/10.34194/geusb.v57.8374 http://www.geusbulletin.org/ voutchkova et al. 2024: geus bulletin 57. 8374. https://doi.org/10.34194/geusb.v57.8374 6 of 10 geusbulletin.org with °dhwsa and n. figure s8 (in supplementary file s1) compares the arithmetic and volume-weighted means. estimating annual production volumes field ‘estimmaengde’ contains an estimated annual groundwater abstraction for the waterworks (anlaegid) or, if not available, for the groundwater wells belonging to the specific waterworks. when annual groundwater abstractions were reported for both waterworks and their well screens, the largest value was used. zero volumes were treated as reporting errors and substituted with ‘missing value (na)’. it was assumed that if there has been a compliance sampling in a specific year, dw must have also been produced – however, the volume was wrongly reported to jupiter. time series for volume (annual resolution) were arranged by waterworks id and year. for each waterworks, the missing values were filled using the ‘downup’ method, which first looked for values down, then up in the specific time series (wickham et al. 2023b). filling missing values in the time series was only possible when there was at least one value in the entire period. for the waterworks without any reported volumes in the period, and for wsas with multiple waterworks, a mean volume was calculated based on the information from the other waterworks in the wsa (max. n = 7) and assigned to the waterworks with a missing volume for the specific year. this resulted in either underor over-estimating the actual volumes, but it was preferred over excluding waterworks from the volume-weighted mean calculation, eq. (4), where it would have otherwise left a missing value. it was important to retain information on dw hardness at a waterworks, even when there was no information on the volume. even after that, there were waterworks with missing volumes for some years, so another round of gap filling was done at the waterworks level following the same procedure. to summarise, missing annual volumes were imputed by (1) time-series gap-filling at the waterworks level, followed by (2) spatial averaging at wsas for the specific year, followed by (3) time-series gap-filling at the waterworks level. this left 74 missing values for 7 wsas, so arithmetic mean hardness at the wsas was used there instead. estimating the current dw hardness at the wsa the final data product (data set 3) contains the latest dw hardness at the wsa based on the data from data set 2. it was assumed that the current dw hardness at a wsa is equal to the latest estimated annual hardness at the wsa. this was done to maximise the spatial data coverage while still showing the most recent available data. the latest annual hardness was linked to the current wsa polygons (end = 2019). software all described procedures were implemented in r (v. 4.2.1; r core team 2022), using the r packages: tidyr (v. 1.3.0; wickham et al. 2023b), dplyr (v. 1.1.0; wickham et al. 2023a), data.table (v. 1.14.8; dowle & srinivasan 2023). all graphs were produced with ggplot2 (v. 3.4.1; wickham 2016). the maps were made with qgis (v. 3.22.10; qgis development team 2021). data description and main features data set 1 data set 1 (fig. 2) contains information about the annual dw hardness at the public waterworks in denmark (n = 3963) between 1905 and 2023 (see supplementary file s2 for a description of the contents). reporting of dw quality by the danish counties became mandatory in 1986 and digitising of old laboratory records was not uniform throughout denmark. this is reflected in the data availability (fig. 2). data set 1 can be used to study the temporal changes in dw hardness at individual waterworks. overall, there has been a decrease in the proportion of public waterworks supplying extremely hard, very hard and hard water, and an increase in the proportion of medium hard and fairly hard dw (fig.  2). following the de-regulation (see section ‘background information’), there were fig. 2 drinking water hardness at public waterworks in denmark. vertical dashed line shows when ca, mg and total hardness were de-regulated (see text for details). x-axis is limited to 1950–2023 as there are very few waterworks with data in the period 1900–1950. sampling and reporting of data for 2023 was incomplete because the data were downloaded in september 2023. data available in supplementary files s4 (annual_hardness_public_waterworks.csv). 0.00 0.25 0.50 0.75 1.00 0 500 1000 1500 2000 sampling year p ro po rti on o f w at er w or ks p er h ar dn es s cl as s w aterw orks (n) waterworks (n) 1950 1975 2000 2025 hardness very soft soft medium hard fairly hard hard very hard extremely hard https://doi.org/10.34194/geusb.v57.8374 http://www.geusbulletin.org/ voutchkova et al. 2024: geus bulletin 57. 8374. https://doi.org/10.34194/geusb.v57.8374 7 of 10 geusbulletin.org on average 1540 waterworks with dw hardness data (2018–2022), which is c. 200 waterworks less than in the period 2013–2017 (n = 1765). data set 2 data set 2 (fig. 3) contains annual dw hardness at the wsas of public waterworks in denmark (n = 3571) between 1978 and 2023 (see supplementary file s2 for a description of the contents). the period is limited due to wsa availability from 1978. data set 2 can be used in exposure assessments by spatially linking it to geocoded households to estimate the proportion of households in each hardness class and how these proportions have changed through time. figure 3 shows change in wsas per hardness class. note that the wsas are of different sizes, and the number of households within wsas is not equal. in the 5-year period following the de-regulation (2018–2022), there were on average 1490 wsas with hardness estimates, while in the preceding period (2013–2017) there were 1779 wsas. however, there were still wsas with missing annual dw hardness estimates in the period 2013–2017. fig. 3 drinking water hardness at the water supply area (wsa) of public waterworks (1978–2023). the vertical dashed line shows when ca, mg and total hardness were de-regulated (see text for details). data available in supplementary files s4 (annual_hardness_wsa.csv). 0.00 0.25 0.50 0.75 1.00 0 500 1000 1500 2000 1980 2010 sampling year p ro po rti on o f w s a p er h ar dn es s cl as s w s a (n) hardness very soft soft medium hard fairly hard hard very hard extremely hard wsa (n) 1990 2000 2020 fig. 4 drinking water hardness at the water supply area (wsa) of public waterworks in denmark, according to the most recently available data between 1980 and 2023. data extracted from the jupiter database 6 september 2023. data available in supplementary file s6. drinking water hardness * very soft soft medium hard fairly hard hard very hard extremely hard wsa without data outside wsa 0 25 50 75 100 km n * latest data available for 1980–2023 (fig. 5) data extraction from jupiter database: 6 sep 2023 https://doi.org/10.34194/geusb.v57.8374 http://www.geusbulletin.org/ voutchkova et al. 2024: geus bulletin 57. 8374. https://doi.org/10.34194/geusb.v57.8374 8 of 10 geusbulletin.org data set 3 data set 3 (fig. 4) contains the most recent available information about dw hardness at the wsas of public waterworks (see supplementary file s2 for a description of the contents). these data can be used for visualisation purposes in gis or in online viewers. data set 3 differs from data set 2 because it only includes one value per wsa – the most recent one – while data set 3 includes all annual data at the wsa. figure 5 shows the corresponding year for each wsa. the most recent annual hardness estimates for 66% (n = 1677) of wsas are for 2022 or 2023, and 34% (n = 883) for the period between 1980 fig. 5 most recently available drinking water (dw) hardness estimates at the water supply area (wsa) of public waterworks in denmark. (a) map of data distribution in denmark. (b) histogram showing the number of wsas and the latest year when dw hardness data are available. data available in supplementary file s7. latest available dw data 1980–1990 1990–2000 2000–2010 2010–2020 2020–2023 wsa without data outside wsa 0 25 50 75 100 km (a) 0 250 500 750 1980 2000 2020 latest year with dw hardness data w s a (n ) (b) data extraction from jupiter database: 6 september 2023 n 1990 2010 https://doi.org/10.34194/geusb.v57.8374 http://www.geusbulletin.org/ voutchkova et al. 2024: geus bulletin 57. 8374. https://doi.org/10.34194/geusb.v57.8374 9 of 10 geusbulletin.org and 2021. the distribution of wsas in each hardness class is as follows: very soft 0.6%, soft 9.5%, medium hard 25.7%, fairly hard 43.1%, hard 18.8%, very hard 2.1% and extremely hard 0.1%. this shows that most danish wsas (by number) are supplied with fairly hard to hard water. there is substantial uncertainty in the abstraction volumes due to the gap-filling procedures used, but roughly 65% of the abstracted water (by volume) falls into fairly hard to hard categories. more meaningful results could be obtained if the proportion of supplied households were used instead, but that was out of the scope of this article. summary dw hardness may vary locally and temporally. the three aggregated data sets reported here are meant to be used for regional or national scale overviews and assessments. the chosen methodology is currently the optimal way to make an overview of dw hardness at different scales in denmark. however, due to the limitations listed here and in schullehner (2022) concerning the actual wsas, dw quality and abstraction volumes, we recommend contacting the water utility companies if current information is needed about the hardness of the tap water for a specific consumer. acknowledgements the authors would like to acknowledge the feedback about de-regulation of ca, mg and total hardness, which we received at the danish water forum 2024. additional information funding statement this study was supported by internal funding from the geological survey of denmark and greenland (geus). authors’ contributions ddv: data curation, formal analysis, software, methodology, visualization, writing – original draft, writing – review & editing ctt: data curation, methodology, software, writing – review & editing laol: data curation, software, writing – review & editing nc: formal analysis, software, writing – review & editing lts: funding acquisition, methodology, validation, writing – review & editing bpj: funding acquisition, methodology, project administration, writing – review & editing bgh: methodology, validation, writing – review & editing competing interests the authors declare no competing interests. additional files a collection of seven files and 3 compressed data set files are available to download with this article at https://doi.org/10.22008/fk2/ nlx5sx. supplementary file s1 – supplementary_materials.docx. supplementary file s2 – readme.txt. supplementary files s3 (set of files containing input data: input1.csv, input2.csv). supplementary files s4 (set of csv files containing data sets 1, 2, and 3: annual_hardness_public_waterworks.csv, annual_hardness_wsa. csv, latest_hardness_wsa.csv). supplementary files s5 (set of files containing shape file for data set 1). supplementary files s6 (set of files containing shape file for data set 2). supplementary files s7 (set of files containing shape file for data set 3). data set disclaimer geus shall not be held liable for any loss or damage, whether direct or indirect, because of the use of the hardness data sets. it is recommended that users obtain additional information from the relevant utilities before taking decisions regarding water hardness. references council of the european union. 2020: directive (eu) 2020/2184 of the european parliament and of the council of 16 december 2020 on the quality of water intended for human consumption. official journal of the european union l 435/1, 62. dowle, m. & srinivasan, a. 2023: data.table: extension of ‘data.frame’. https://cran.r-project.org/package=data.table (accessed february 2024). geological survey of denmark and greenland (geus). 2010: drikkevandets hårdhed i danmark. 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qgis association; https://qgis.org/ (accessed august 2024). https://doi.org/10.34194/geusb.v57.8374 http://www.geusbulletin.org/ https://doi.org/10.22008/fk2/nlx5sx https://doi.org/10.22008/fk2/nlx5sx https://cran.r-project.org/package=data.table https://doi.org/10.22008/fk2/5ozv7w https://doi.org/10.22008/fk2/5ozv7w https://data.geus.dk/tabellerkoder/koder.html?codetype=744 https://data.geus.dk/tabellerkoder/koder.html?codetype=744 https://www.miljoeogressourcer.dk/filer/lix/5033/geovejledning_nr._6__2018__978-87-7871-493-0_.pdf https://www.miljoeogressourcer.dk/filer/lix/5033/geovejledning_nr._6__2018__978-87-7871-493-0_.pdf https://doi.org/10.34194/geusb.v23.4842 https://www.retsinformation.dk/eli/lta/1988/515 https://www.retsinformation.dk/eli/lta/1988/515 https://www.retsinformation.dk/eli/lta/2001/871 https://www.retsinformation.dk/eli/lta/2001/871 https://www.retsinformation.dk/eli/lta/2017/1147 https://www.retsinformation.dk/eli/lta/2017/1147 https://www.retsinformation.dk/eli/lta/2022/602#id314a3088-4186-4900-ad88-ba56b5967af5 https://www.retsinformation.dk/eli/lta/2022/602#id314a3088-4186-4900-ad88-ba56b5967af5 https://www.retsinformation.dk/eli/lta/2023/1023 https://www.retsinformation.dk/eli/lta/2023/1023 https://www.retsinformation.dk/eli/lta/2023/529 https://cdnmedia.eurofins.com/microsites/media/2778/m069-metaller-grv-drv-02.pdf https://cdnmedia.eurofins.com/microsites/media/2778/m069-metaller-grv-drv-02.pdf https://davidmathlogic.com/colorblind/ https://qgis.org/ voutchkova et al. 2024: geus bulletin 57. 8374. https://doi.org/10.34194/geusb.v57.8374 10 of 10 geusbulletin.org r core team. 2022: r: a language and environment for statistical computing. r foundation for statistical computing. vienna, austria. https://www.r-project.org/ rygaard, m. & albrechtsen, h.-j. 2020: blødgøring af drikkevand i aarhus – forventede konsekvenser. 24 pp. dtu, institut for vand og miljøteknologi. copenhagen. https://orbit.dtu.dk/files/240820763/ bl_dg_ring_af_drikkevand_i_aarhus_210104.pdf (accessed january 2024). schullehner, j. 2022: danish water supply areas and their links to water production facilities: an open-access data set. geus bulletin 49, 8319. https://doi.org/10.34194/geusb.v49.8319 voutchkova, d.d., schullehner, j., skaarup, c., wodschow, k., ersbøll, a.k. & hansen, b. 2021: estimating pesticides in public drinking water at the household level in denmark. geus bulletin 47, 6090. https://doi. org/10.34194/geusb.v47.6090 wickham, h. 2016: ggplot2: elegant graphics for data analysis. 182 pp. springer-verlag new york. https://ggplot2.tidyverse.org. wickham, h., françois, r., henry, l., müller, k. & vaughan, d. 2023a: dplyr: a grammar of data manipulation. https://cran.r-project.org/ package=dplyr (accessed august 2024). wickham, h., vaughan, d. & girlich, m. 2023b: tidyr: tidy messy data, r package. https://cran.r-project.org/package=tidyr (accessed august 2024). https://doi.org/10.34194/geusb.v57.8374 http://www.geusbulletin.org/ https://www.r-project.org/ https://orbit.dtu.dk/files/240820763/bl_dg_ring_af_drikkevand_i_aarhus_210104.pdf https://orbit.dtu.dk/files/240820763/bl_dg_ring_af_drikkevand_i_aarhus_210104.pdf https://doi.org/10.34194/geusb.v49.8319 https://doi.org/10.34194/geusb.v47.6090 https://doi.org/10.34194/geusb.v47.6090 https://ggplot2.tidyverse.org https://cran.r-project.org/package=dplyr https://cran.r-project.org/package=dplyr https://cran.r-project.org/package=tidyr open access nationwide data sets for drinking water hardness at public waterworks and their water supply areas in denmark data collection background information data sources data preparation statistical analysis and data processing semi-automated outlier detection procedure dw hardness calculation and classification data aggregation at the waterworks level link between waterworks and wsas estimating the dw hardness at wsas estimating annual production volumes estimating the current dw hardness at the wsa software data description and main features data set 1 data set 2 data set 3 summary acknowledgements additional information funding statement authors’ contributions competing interests additional files data set disclaimer references figures fig. 1 flowchart summarising the data preparation and statistical analyses used here. qa/qc: quality assurance and quality control. wsa: water supply area. for the contents of data sets 1, 2, and 3, see section ‘data description and main features’. hardness calculated by eq. (3). field names: anlaegid: unique identification number (id) of the waterworks. plantid: waterworks id used in schullehner (2022), same as anlaegid. proveid: sample id. proeveaar: sample year. wsaid: id of the wsa (schullehner 2022). full list of field names in supplementary files s1, s2. fig. 2 drinking water hardness at public waterworks in denmark. vertical dashed line shows when ca, mg and total hardness were de-regulated (see text for details). x-axis is limited to 1950–2023 as there are very few waterworks with data in the period 1900–1950. sampling and reporting of data for 2023 was incomplete because the data were downloaded in september 2023. data available in supplementary files s4 (annual_hardness_public_waterworks.csv). fig. 3 drinking water hardness at the water supply area (wsa) of public waterworks (1978–2023). the vertical dashed line shows when ca, mg and total hardness were de-regulated (see text for details). data available in supplementary files s4 (annual_hardness_wsa.csv). fig. 4 drinking water hardness at the water supply area (wsa) of public waterworks in denmark, according to the most recently available data between 1980 and 2023. data extracted from the jupiter database 6 september 2023. data available in supplementary file s6. fig. 5 most recently available drinking water (dw) hardness estimates at the water supply area (wsa) of public waterworks in denmark. (a) map of data distribution in denmark. (b) histogram showing the number of wsas and the latest year when dw hardness data are available. data available in supplementary file s7. table table 1 drinking water (dw) hardness classification and recommended colour palette to represent each class. research article andrews et al. 2021: geus bulletin 47. 6523. https://doi.org/10.34194/geusb.v47.6523 1 of 14 the permian to cretaceous succession at permpasset, wollaston forland: the northernmost permian and triassic in north–east greenland steven d. andrews*1 , henrik nøhr-hansen2 , pierpaolo guarnieri2 , karen dybkjær2 , sofie lindström2 , peter alsen2 1university of the highlands and islands, inverness, uk, 2geological survey of denmark and greenland (geus), copenhagen, denmark abstract permian to triassic outcrops in east greenland diminish significantly northwards. understanding the northward extent, and nature, of the permian and triassic successions has implications for regional palaeogeographic reconstructions and exploration in adjacent offshore basins. examining the structural relationships between the basement, permian, triassic, jurassic and cretaceous successions can further our understanding of the tectonic evolution of the region. here, we describe a hitherto overlooked section through the permian to cretaceous from central wollaston forland and consider its structural context. the western side of permpasset forms the upthrown eroded crest of a horst block, which provides exposure of the earliest stratigraphic intervals in the region. the fractured caledonian basement is overlain by evaporitic marine limestone facies of the karstryggen formation, which are succeeded by shallow marine sandstones assigned to the schuchert dal formation, both upper permian. the overlying unit records a period of fluvial deposition and is not possible to date. however, an early to middle triassic age (pingo dal group) seems most likely, given regional eustatic considerations. this is, therefore, the most northerly record of triassic strata in north–east greenland. west of the horst structure, fine-grained sandstones and bioturbated siltstones of the jurassic (oxfordian) jakobsstigen formation are recorded. these were downfaulted prior to a prolonged hiatus after which both the triassic and jurassic strata were draped by cretaceous shales of the fosdalen formation. the cretaceous succession is overlain by a thick basalt pile of eocene age, heralding the opening of the north-east atlantic. glendonites overlie oxfordian siltstones at the base of the middle albian fosdalen formation. these were likely winnowed from slightly older cretaceous strata and overlie the hiatus surface between the jurassic and cretaceous. this is the first record of glendonites from the cretaceous of east greenland and they are interpreted to record the circum–arctic late aptian – early albian cooling event. *correspondence: steven.andrews914@ gmail.com received: 12 jan 2021 accepted: 20 may 2021 published: 23 july 2021 keywords: cretaceous, east greenland, glendonites, permian, triassic abbreviations: mierl: ministry of industry, energy, research and labour geus: geological survey of denmark and greenland geus bulletin is an open access, peerreviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: mette olivarius (geus, denmark) reviewed by: ashton embry (geological survey of canada, canada), tore grane klausen (petrolia noco as, norway) funding: see page 13 competing interests: none declared additional files: none provided 1 introduction and geological background the permian and triassic have received much attention where spectacularly exposed in jameson land, scoresby land, traill ø, geographical society ø, hold with hope and clavering ø (clemmensen 1980a, 1980b; surlyk et al. 1986; christiansen et al. 1993; stemmerik 2001; andrews et al. 2019, https://doi.org/10.34194/geusb.v47.6523 https://orcid.org/0000-0002-0418-0809 https://orcid.org/0000-0002-9291-8104 https://orcid.org/0000-0002-7509-0050 https://orcid.org/0000-0002-8420-3379 https://orcid.org/0000-0001-8278-1055 https://orcid.org/0000-0001-6218-9054 mailto:steven.andrews914@gmail.com mailto:steven.andrews914@gmail.com andrews et al. 2021: geus bulletin 47. 6523. https://doi.org/10.34194/geusb.v47.6523 2 of 14 www.geusbul let in.org 2020a, 2020b; clemmensen et al. 2020). however, those sections lying further north on wollaston forland have seen little, if any, concerted study (fig. 1). the nature of the northward continuation of the permian and triassic successions has important implications for palaeogeographic reconstructions and our understanding of adjacent offshore basins, and as such these outcrops have an increased significance. maync (1942) described the permian succession of wollaston forland, but with the advancement of our understanding of the regional setting and stratigraphy, a number of questions existed about this remote and seldom visited outcrop. of particular interest is the enigmatic clastic succession that overlies the permian limestones. furthermore, examining the structural relationships among the basement, permian, triassic, jurassic and cretaceous successions can further our understanding of the tectonic evolution of the region. this work was undertaken as part of a larger mierl (ministry of industry, energy, research and labour) and geus (geological survey of denmark and greenland)-led regional scale mapping project, which included examination of the full permian to eocene succession and the investigation of the associated structural elements across the wollaston forland. as part of this project, time was spent examining the permian to paleocene–eocene stratigraphy and associated structure of permpasset, central wollaston forland. the results of these investigations are presented here. this includes an assessment of the stratigraphy, which is thought to include permian, triassic, jurassic cretaceous and paleocene–eocene strata, capped by a thick plateau basalt succession. the exposure of the deeper stratigraphic levels recorded in permpasset occurs due to the presence of the n–s-orientated permpasset fault, which downthrows to the east and exposes the eroded crest of the permpasset horst block, including the succession described here. the western margin of the horst is defined by an ne–sw-orientated fault, which converges on the permpasset fault to the north. a further e–w-orientated fault divides the horst but appears to have limited displacement. these structures and their temporal relationships are described in detail in the next sections. 2 methodology geological mapping was undertaken by foot traverse supported by 3d-photogeology in the geus photogrammetry lab (sørensen & dueholm 2018; sørensen & guarnieri 2018) based on new oblique photos acquired in 2017 (olsen & jakobsen 2018). sedimentological logging was undertaken at 1:40 scale where exposures were of good quality and at 1:2000 scale where exposures were poor. samples of fine-grained lithologies were collected throughout the successions for palynological analysis. macropalaeontological specimens were also collected for identification. palynological samples were prepared using standard laboratory techniques (poulsen et al. 1990) before being counted and identified. 3 sedimentology and stratigraphy permpasset provides the most easterly exposure of palaeozoic strata in wollaston forland. the succession recorded occurs as a small area of outcrop on the west side of the valley at the apex of the pass. the permpasset horst is defined by the permpasset fault to the east and in the west by an ne–sw-aligned fault (figs 1, 2a). sections were logged through the eroded horst crest and to the north-west in the hanging wall of the ne– sw-trending fault. these are described in turn below. 3.1 permpasset horst description. the lowermost unit is 36 m thick (fig. 3) and rests on heavily fractured caledonian basement mostly comprising granite and quartzite. the initial 15 m is poorly exposed, but the likely lithology is indicated by large blocks of shell-rich limestone, which are unlikely to have travelled far. the shell material includes abundant productid brachiopods and crinoids, which vary between different beds from heavily fragmented to almost complete. overlying this, initially poorly exposed section is a prominent cliff-forming succession. the base of the cliff comprises brecciated limestones, which display rounded clasts with minor intercalations of laminated limestones (3 m). these transitions upward to blocky brecciated limestone and increase volumes of gypsum (fig. 2b), which eventually form the dominant lithology (4 m). blocks up to several metres are recorded. an uneven surface is then infilled by bedded selenite (fig.  2c), which forms a unit up to 1 m thick before being overlain by a further 4 m of poorly exposed brecciated limestone and gypsum, again displaying an uneven upper surface. these are overlain by 8.5 m of laminated limestone devoid of shell material (fig. 2d). minor undulatory lamination is recognised in the basal 2 m (fig. 3). the middle unit reaches a thickness of 23 m and comprises shell-rich, fine-grained, buff-coloured sandstones and grey limestones intercalated with coarsely laminated grey limestones forming beds between 0.1 and 0.3 m thick. the shell material is again dominated by productid brachiopods with beds often containing a mixture of both heavily fragmented material as well as complete valves (fig. 4a). articulated specimens are rare. horizontal lamination predominates, and some beds display erosive bases. overall, an upward increase in both the sandstone component and shell material occurs in this unit (fig. 3). minor bioturbation is also recognised towards the unit top. https://doi.org/10.34194/geusb.v47.6523 www.geusbulletin.org� andrews et al. 2021: geus bulletin 47. 6523. https://doi.org/10.34194/geusb.v47.6523 3 of 14 www.geusbul let in.org fig. 1 geological map of wollaston forland (modified after escher 2001). the black-lined box marks the position of the map illustrated in fig. 7. the inset map indicates the position of wollaston forland (wf) on the east greenland coast as well as other regions mentioned in the text: cø: clavering ø. hwh: hold with hope. gsø: geographical society ø. tl: traill ø. sl: scoresby land. jl: jameson land. 19°w19°30'w20°w 74 °4 0' n 74 °3 0' n 74 °2 0' n 0 2,5 5 7,5 10 km legend ice and perennial snow sea quaternary tertiary dykes and sills plateau basalts sandstones cretaceous stratumbjerg-fosdalen fms palnatokes bjerg fm lindemans bugt fm jurassic jakobsstigen fm pelion fm permian foldvik creek group caledonian basement quartzites and migmatites permpasset jurakløft 20°0'0"w25°0'0"w 75 °0 '0 "n 74 °0 '0 "n 73 °0 '0 "n 72 °0 '0 "n 71 °0 '0 "n cø wf sl hwh gsø tø jl / triassic pingo dal gp https://doi.org/10.34194/geusb.v47.6523 www.geusbulletin.org� andrews et al. 2021: geus bulletin 47. 6523. https://doi.org/10.34194/geusb.v47.6523 4 of 14 www.geusbul let in.org 0.3 m 0.5 m b c d basement p tr? jr k k basalt k p tr? stratigraphic boundary fault logged section a fig. 3 fig. 6 north fig. 2 overview of the permpasset outcrop and photographs illustrating the key lithologies of the permian succession. a: view west across permpasset illustrating the stratigraphic and structural relationships and the position of the logged sections provided in figs 3 and 6. p: permian. tr: triassic. jr: jurassic. k: cretaceous. b: brecciated limestone and gypsum, including progressively larger blocks upwards (karstryggen formation). c: bedded selenite overlying the first brecciated interval (karstryggen formation). d: bedded limestones recording a deepening and freshening of marine conditions, perhaps forming a correlative of the wegener halvø formation. the base of the topmost unit is marked by a distinct coarsening with mediumto coarse-grained pebbly sandstones recognised (fig. 3). trough cross-bedding and planar cross-bedding also become common (fig.  4b), forming sets up to 0.4 m thick and indicating a broadly southerly directed flow. minor bioturbation is recognised, but there is a complete lack of shell material. the lowermost 3 m of sandstones is succeeded by a poorly exposed recessive section of around 7 m comprising red to orange coloured fine-grained/muddy material. this is overlain by a further 23 m of fineto medium-grained, feldspathic, buff-coloured sandstones, again displaying abundant trough cross-bedding. erosive bed bases, often defined by pebble and mudstone rip-up lags, are also common. despite the pale appearance of these uppermost sandstones, when a fresh surface is exposed, a much darker colouration is noted. https://doi.org/10.34194/geusb.v47.6523 www.geusbulletin.org� andrews et al. 2021: geus bulletin 47. 6523. https://doi.org/10.34194/geusb.v47.6523 5 of 14 www.geusbul let in.org cretaceous (fosdalen fm.) triassic? (pingo dal gp.) schuchert dal fm. karstryggen fm. basement c si vf s fs m s cs vc s g p co 10 0 m 20 30 40 50 60 70 80 90 100 black staining throughout red recessive unit laminated limestone stromatolites? stromatolites? 548121 548122 548123 548124 548125 548126&548127 548128 548129 548130 548131 548132 selenite gypsum permpasset central tra ns gr es sio n key horizontal lamination planar cross-bedding bioturbationtrough cross-bedding convolute lamination pebbles evaporites (gypsum) limestone selenite bivalves evaporite sandstone siltstone basement regression fig. 3 logged section from the central outcrops of the permpasset exposures. the location of the section is indicated in fig. 2a. sample numbers are included adjacent to the log. https://doi.org/10.34194/geusb.v47.6523 www.geusbulletin.org� andrews et al. 2021: geus bulletin 47. 6523. https://doi.org/10.34194/geusb.v47.6523 6 of 14 www.geusbul let in.org 0.5 m 0.1 m 0.5 m 10 mm 10 mm a b c d e f g h 20 mm 0.1 m10 mm fig. 4 photographs illustrating the key lithologies of the uppermost permian, and the triassic, jurassic and cretaceous stratigraphy. a: interbedded shelly limestones and sandstones forming the top of the permian succession (schuchert dal formation). b: buff-coloured planar and trough cross-bedded sandstones with intraclast lags commonly present at bed bases (triassic pingo dal group?). c: pebbly sandstone with abundant rounded quartz clasts and occasional sandstone clasts forming the base of the jurassic (jakobsstigen formation). d and e: different glendonite forms from the base of the cretaceous succession. f: graded lamination and contorted intraclasts within the siltstones of the fosdalen formation. g: current ripples with possible signs of storm reworking in sandstone beds weathering out of the fosdalen formation. h: flute marks and thalassinoides ichnofossils preserved on bed bases in the fosdalen formation. https://doi.org/10.34194/geusb.v47.6523 www.geusbulletin.org� andrews et al. 2021: geus bulletin 47. 6523. https://doi.org/10.34194/geusb.v47.6523 7 of 14 www.geusbul let in.org this succession is capped by a thin (2 m) unit of very coarse, quartz rich, pebbly sandstones, above which the slopes are recessive, potentially indicative of finer-grained facies. interpretation and stratigraphic assignation. the brecciated and evaporitic nature of the lowermost unit most closely resembles the karstryggen formation as originally described from the jameson land region (surlyk et al. 1986). this is also consistent with the abundant late permian marine fauna. the features described are interpreted as evidence of a marine environment where exposure led to the development of brecciation through karstification during subaerial exposure. the first brecciated interval is overlain by bedded selenite, suggesting not only renewed flooding of the region but also the development of strongly evaporative conditions, most likely reflecting a highly restricted setting. the second brecciated interval is topped by laminated limestones, which suggest a return to deeper water, non-evaporative conditions. this is consistent with the trends observed elsewhere in east greenland, where the karstryggen formation is topped by the wegener halvø formation reefal limestones, and the black shales of the ravnefjeld formation (surlyk et al. 1986; christiansen et al. 1993; stemmerik 2001). the laminated limestones described here, although likely to correlate with the deepening conditions recorded by these units, are considered too thinly developed to define as a separate formation. the middle unit is interpreted to reflect a gradual return to shallower water conditions, initially with the introduction of occasional shell-rich sandy beds before facies become dominated by horizontally laminated sandstones, which, alongside the fragmentation of shell material, provide evidence for high-energy conditions, most likely reflecting a shallow marine environment. the sandy nature of this unit and the regressive signature are most similar to the schuchert dal formation described to the south, which occurs as both shallow marine and turbiditic facies (surlyk et al. 1986; kreiner-møller & stemmerik 2001). the topmost unit is marked by its complete absence of shell material. the presence of planar and trough cross-bedding, alongside erosive bed bases containing mudstone rip-up clasts and the red to orange colouration of the poorly exposed intervening finer-grained units is suggestive of a continental, fluvial depositional environment. the age of this 32 m thick succession is problematic with permian, triassic or even jurassic ages being possible. no late permian fluvial deposits have been reported from east greenland. furthermore, this period is characterised by falling sea levels and the development of a hiatus, particularly outside of basinal areas (surlyk et al. 1986). therefore, the preservation of fluvial deposits at this time seems unlikely. the early triassic is marked by a major transgression, leading to the widespread deposition of the marine wordie creek formation (seidler et al. 2004). no evidence is found in the uppermost unit described here for marine conditions. to the south, on clavering ø, the wordie creek formation comprises pebbly buff sandstones and green-grey shales. although the facies described here do resemble these buff sandstones, they lack the hummocky cross stratification noted in the clavering ø examples, and the intervening finer-grained facies appear to be red-orange in colour, favouring a continental origin. the next candidate would be the continental, early to middle triassic pingo dal group. the depositional environment and stratigraphic position would favour the interpretation of this unit as the pingo dal group, but one further option requires consideration, the jurassic, and more specifically, the continental bastians dal or bristol elv formations. however, both these units are characterised by abundant organic/woody debris and their high compositional maturity, whereas the topmost unit described here is devoid of organic remains and also appears to be highly feldspathic. therefore, the simplest interpretation that can be made would place this unit as the early to middle triassic pingo dal group. the age of the pebbly sandstone, which caps the exposure, is unknown, but it is lithologically distinct from the underlying sandstones. the overlying recessive slopes most likely represent fine-grained facies of the mid-cretaceous strata, which blanket the region. therefore, the pebbly sandstone could be interpreted as representing condensed deposits, which overlie the unconformity surface prior to the onset of mid-cretaceous deposition. evidence for reworked triassic in the region. although the unit described above appears isolated, with respect to the regional distribution of triassic strata, reworked triassic palynomorphs have been recorded in outcrop samples of the jurassic pelion and payer dal formations in the jurakløft area, northern wollaston forland. these formations are dated by ammonites and dinoflagellate cysts to the bathonian–callovian and oxfordian, respectively (this study). the reworked palynomorphs are generally darker in colour than the jurassic dinocysts but differ little in appearance from the jurassic spores and pollen. identified triassic pollen includes protodiploxypinus gracilis, podosporites amicus, triadispora  sp. and aulisporites astigmosus (fig. 5). the recovered triassic spore/pollen assemblage is believed to indicate reworking of late triassic, carnian, strata. this is based primarily on the biostratigraphic distribution of aulisporites astigmosus, which is more or less confined to the carnian in the boreal area (vigran et al. 2014; paterson & mangerud 2015). on svalbard and in the barents sea, protodiploxypinus sp. has an acme in the latest carnian, https://doi.org/10.34194/geusb.v47.6523 www.geusbulletin.org� andrews et al. 2021: geus bulletin 47. 6523. https://doi.org/10.34194/geusb.v47.6523 8 of 14 www.geusbul let in.org where it occurs with sporadic aulisporites astigmosus (paterson & mangerud 2015, 2020). in the jurakløft assemblages, a. astigmosus is one of the most common reworked palynomorphs. this may indicate that strata equivalent of the carnian pluvial episodes (e.g. dal corso et al. 2020) were once present in the area. farther south, on jameson land, palynological assemblages interpreted to represent the carnian pluvial episode appear to be lacking aulisporites astigmosus (andrews et al. 2014). however, further north on spitsbergen a. astigmosus sometimes makes up >20% of the late julian assemblages from tschermakfjellet formation (mueller et al. 2016). 3.2 north west of the permpasset horst description. this section lies to the north-west of the horst where strata have been downfaulted to the west (fig. 2a) and comprise three distinct units. the base of the section comprises a very coarse-grained, quartzrich, pebbly sandstone containing pebbles up to 5 cm (fig. 4c). pebble lithologies are dominated by quartz but minor schist/gneiss and sandstone clasts are recognised. exposure is poor above this coarse-grained unit, but the recessive slopes are composed of medium to fine-grained sands, which continue upwards for 25 m (fig. 6). these are overlain by interbedded light grey silty mudstones and fine-grained, poorly consolidated sandstones, which occasionally contain small pebbles (2–5 mm). bioturbation and bivalve moulds are recognised within this unit as well as occasional large wood fragments. upwards, sandstone beds become less common but bioturbation appears to increase in intensity producing a mottled appearance to the silty mudstones. coarsening-upward cycles (10–25 m thick) are recognised within this unit, with each cycle capped by siltstones to sandy siltstones. large carbonate concretions (0.5–1 m) are also recorded throughout, but with a tendency to concentrate towards the top of the cycles described. the concretions themselves often preserve original sedimentary features, which existed within the silts. a distinct bench overlain by a 10 m thick mudstone marks the top of this unit. the topmost unit is marked by very dark to black mudstones at its base. these are initially bioturbated, but upwards bioturbation dies out. the base of this unit is further marked by the presence of glendonites, which are common in a single interval (figs 4d, e). belemnite fragments are also recorded at this horizon. the dark colouration continues for 55 m. in this interval, bioturbation is not common, but siderite nodules and carbonate nodules displaying cone-in-cone structure are noted. rare stream-worn sections reveal 1–10 mm scale lamination displaying grading and containing 1–4 mm deformed mudstone rip-up clasts at laminae bases (fig. 4f). above these dark-coloured mudstones, a general lightening of tone, accompanied by increased silt content and the return of bioturbation, is noted. this forms a further 87 m of section before the base of the overlying basalts is reached (fig. 6). further sedimentological detail for this succession can be gleaned from a laterally correlative section examined to the northwest. here, the bioturbated siltstones contain better defined coarsening-upward cycles on a 10–20 m scale (fig. 6). these cycles are often topped with very fineto fine-grained sandstone beds up to 0.1 m in thickness. bed bases display flute marks (fig. 4h) and occasional load structures. internally, current rippling and possible hummocky cross-stratification (fig. 4g) are noted alongside bioturbation, including well-developed thalassinoides (fig. 4h). the base of the basalts, which overlie the sedimentary succession, is uneven, and in places, large grey carbonate nodules and coalified wood fragments appear to be embedded within a claystone matrix (paleocene– eocene?), which is, in turn, overlain by autobreccia, which appears to incorporate some of the underlying sediment. interpretation and stratigraphic assignation. the presence of bivalve casts, abundant bioturbation as well as the presence of acritarchs and dinocysts recognised during micropalaeotological analysis favour a marine environment for the lowermost unit of this section. a callovian age has been ascertained from the top of this fig. 5 examples of reworked triassic pollen. taxon name is followed by sample and slide number and england finder coordinates. the scale bar is 20 μm. a: aulisporites astigmosus, 589748:4, u58/3. b: protodiploxypinus gracilis, 589757:5, h18/1. c: lunatisporites sp., 589754:5, h53/2. d: protodiploxypinus sp. cf. p. gracilis, 589757:4, s43/1. https://doi.org/10.34194/geusb.v47.6523 www.geusbulletin.org� andrews et al. 2021: geus bulletin 47. 6523. https://doi.org/10.34194/geusb.v47.6523 9 of 14 www.geusbul let in.org 400 500 600 m c si vf s fs m s cs vc s g p co 548145 548146 548147 permpasset nw 548148 548149&548150 548151 ? ? fo sd al en f m . ba sa lts 300 400 500 m c si vf s fs m s cs vc s g p co 548101 548102 548103&546104 548105 548106 548107 548108 548109 (glendonites) 548110 548111 548112 548113 548114 548115 548116 548117 548118 548119&548120 pal/eo permpasset w ja ko bs st ig en f m . fo sd al en f m . ba sa lts key horizontal lamination bioturbation oscillation ripples pebbles bivalves wood sandstone siltstone (plus minor sandstone intercalations) carbonate nodules siderite nodules sill fig. 6 logged sections from the northern portion of permpasset. the location of permpasset w is provided in fig. 2a. pal/eo: paleocene or eocene. sample numbers are included adjacent to the log. https://doi.org/10.34194/geusb.v47.6523 www.geusbulletin.org� andrews et al. 2021: geus bulletin 47. 6523. https://doi.org/10.34194/geusb.v47.6523 10 of 14 www.geusbul let in.org unit based on the recovery of the acritarch veryhachium sortehatense, which has been reported from callovian at hold with hope by piasecki et al. (2004). the presence of sandstone clasts within the basal pebbly sandstone suggests that the permian or triassic sandstones on the adjacent horst may have been exposed and were being reworked at this time. this would also be in accord with the widely recognised middle jurassic rifting and associated erosion of fault block crests (surlyk 1990; surlyk & korstgård 2013). the initial sand-rich strata found at the base of the lower unit gradually give way to intensely bioturbated siltstones (fig. 6). an overall deepening trend is suggested, resulting in both a reduction in sand supply and energy. regular coarsening-upward cycles overprint the overall fining signature. the poorly preserved dinocysts, pareodinia prolongate and nannoceratopsis pellucida, recovered from the very top of this unit suggest a lower to middle oxfordian age. therefore, this unit can be assigned to the jakobsstigen formation. vosgerau et al. (2000) recognised 0.1–19 m alternations between shallow marine and coastal plain conditions in the jakobsstigen formation. the cycles recorded here may be a distal equivalent of these. however, land was likely not too far away, as indicated by the presence of large wood fragments. as age constraints are only available for the top of this unit, it could be speculated that the lowermost, sand-rich portion corresponds to the pelion formation on the basis of its lithological affinity with this callovian unit. however, the poor quality of the outcrop and the lack of biostratigraphic control do not provide confidence in this interpretation. the base of the topmost unit is marked by a distinct darkening in the facies, suggesting a change to poorly oxygenated conditions. this is further evidenced by the absence of bioturbation. also recorded at the contact with the underlying unit is a glendonite-bearing horizon. micropalaeotological analysis of this unit provides a middle albian age, indicated by the presence of the lithosphaeridium arundum and the chichaouadinium vestitum subzones of nøhr-hansen (1993) and nøhr-hansen et al. (2020). however, the most likely date for the glendonites is late aptian – early albian during which time a cooling event is recorded throughout the arctic (rogov et al. 2017). this would, therefore, be the first reported occurrence of glendonites of this age in east greenland. their presence at this stratigraphic boundary, which appears also to represent a significant hiatus, is most easily accounted for if they were winnowed from slightly older cretaceous strata, forming a transgressive lag, prior to the deposition of the middle albian succession. this would explain the juxtaposition of quite different glendonite forms within a single horizon (figs 4d, e). this is consistent with the eastward increase in the erosion of aptian strata across wollaston forland as reported by bjerager et al. (2020). above the dark basal portion of this unit, the presence of intense bioturbation, including thalassinoides, is suggestive of marine conditions. the presence of thin sandstone beds with flute-marked bases and containing current ripples is consistent with turbidite deposition. minor wave reworking may have also occurred during storm activity, as evidenced by the presence of hummocky cross-stratification. graded lamination within the siltstones and micro-scale intraclasts provides evidence for more dilute underflow activity. the cycles recognised, therefore, likely reflect the progradation and retrogradation of more marginal systems during successive transgressive–regressive cycles, which resulted in the advance and retreat of associated turbidite systems. the middle albian age determined for this unit places it within the fosdalen formation. 4 structure of the permpasset horst permpasset is characterised by three sets of normal faults (fig. 7a). cross-cutting relationships demonstrate that the most recent activity occurred along the almost n–s trending permpasset fault (haller 1971; surlyk & korstgård 2013) as indicated by the displacement of the palaeogene basalts, which are down-faulted to the east. the permpasset fault cuts a ne–sw-trending fault. a subsidiary e–w fault is bracketed between these two faults. both the ne–swand the e–w-trending faults post-date the triassic pingo dal group and are sealed by the dark mudstones of the fosdalen formation (fig. 7b). the stratigraphic successions in this small area record two episodes of block tilting. bedding of the permian– triassic strata dips 22–26° toward the west, while the fosdalen mudstones dip 12° to the west. back-tilting of strata shows that the latter tilting most likely relates to the movement along the permpasset fault, which displays more than 400 m offset, while the earlier tilting probably occurred during early cretaceous rifting. jurassic strata display similar dips (23–26° towards the east) to the permian–triassic succession, but as the contact between these units is faulted, their relationship cannot be resolved. according to the data presented here and the cross-cutting relationships, the ne–sw-orientated fault is the only early cretaceous fault as it post-dates oxfordian strata and pre-dates mid–cretaceous mudstones. early cretaceous activity of the permpasset fault was supposed by haller (1971) and surlyk & korstgård (2013). the permpasset horst, as defined here, was interpreted as the footwall crest of a tilted fault block related to the early cretaceous rifting. we cannot exclude cretaceous, or earlier, activity on the permpasset fault, but no direct evidence for earlier movements is available. the geometry of the fault-bounded permapasset horst indicates that it was a structural high located in the central https://doi.org/10.34194/geusb.v47.6523 www.geusbulletin.org� andrews et al. 2021: geus bulletin 47. 6523. https://doi.org/10.34194/geusb.v47.6523 11 of 14 www.geusbul let in.org stratigraphic log !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! !! o o o o o oo o o o o23 26 23 12 14 12 16 22 650 70 0 600 55 0 75 0 45 0 500 35 0 30 0 800 250 85 0 900 950 400 400 450850 500 850 950 600 650 80 0 750 900 800 850 70 0 550 90 0 40 0 32 2622 0 0.5 1 1.5 2 km " " " " " " " " " " " " " " " " " " " " " " " " " " " " " " " " " " " " " " " " permpasset enewsw 1000 800 600 400 200 1000 800 600 400 200 a a' a a' legend boundary o beddingquaternary plateau basalts u. paleoc.-l. eocene fosdalen fm (middle albian) jakobsstigen fm (oxfordian) pingo dal gp (lower-middle triassic) schuchert dal fm (upper permian) karstryggen fm (upper permian) caledonian basement !! unconformity fault, concealed bedding permpasset (a) (b) fig. 7 geological map of the permpasset region and associated cross-section. a: detailed geological map of permpasset from fieldwork and 3d–photogeology. all heights given are metres a.s.l. caledonian migmatites are unconformably overlain by upper permian limestones of the karstryggen formation followed by a sandy succession of the schuchert dal formation (not possible to distinguish in the northern portion of the outcrop due to erosion) and red beds of the pingo dal group. mid–albian mudstones of the fosdalen formation unconformably cover both the pingo dal group and jakobsstigen formation and post-date the ne–sw-trending normal fault. the n–s-trending permpasset fault was active in palaeogene times as shown by almost 400 m offset of plateau basalts to the east and responsible for the westward tilting of the successions. b: wsw–ene geological cross-section of permpasset showing structural relationships and the geometry of the fault-bounded permpasset horst. https://doi.org/10.34194/geusb.v47.6523 www.geusbulletin.org� andrews et al. 2021: geus bulletin 47. 6523. https://doi.org/10.34194/geusb.v47.6523 12 of 14 www.geusbul let in.org part of the wollaston forland basin and associated with first-order n–s-trending faults and segmented by second-order ne–sw faults. the horst was active during early cretaceous rifting and unconformably draped by mudstones during the post-rift thermal subsidence before being uplifted again in palaeogene times, subsequent to the emplacement of the plateau basalts. the horst structure described from geographical society ø by andrews & decou (2019) and andrews et al. (2020a) contains a number of similarities with the permpasset horst and illustrates that structures of this nature may be relatively common. 5 discussion and conclusions the northward extent, and the nature of, palaeozoic and mesozoic facies in east greenland hold important implications for palaeogeographic reconstructions and our understanding of adjacent offshore regions. jurassic and cretaceous strata can be traced northward up the east greenland coast to just north of 76°. however, the nature and the extent of the permian and triassic strata are less well known. the exposures described here provide one of the most northern exposures of permian strata and a candidate for the most northerly triassic strata yet described from east greenland. although the permian succession is much condensed, the karstryggen formation (including a potential correlative of the wegner halvø formation) and the schuchert dal formation are recognised. in both instances, these appear to reflect relatively shallow-water conditions, with the deeper water source rock facies of the ravnefjeld formation being absent. christiansen et al. (1993) noted that only the lower two sequences (karstryggen formation and the wegner halvø/ ravnefjeld formations) described by surlyk et al. (1986) could be traced northwards, with reference to their work on clavering ø, suggesting that a more prolonged period of erosion occurred during the latest permian – earliest triassic in this region. however, the identification of the schuchert dal formation in permpasset provides evidence for the northward continuation of this uppermost permian sequence. therefore, rather than the suggested northward increase in erosion in the latest permian – earliest triassic, the distribution of the schuchert dal formation may be primarily the product of palaeogeographic or palaeoenvironmental controls on its development. certainly, a position within the basin where clastic material is available would have provided a primary control on its development. in terms of the distribution of permian strata on wollaston forland, surlyk & korstgård (2013) suggest a north–eastward thinning and onlap of the caledonian basement. this does not appear to be borne out by geological maps of the region (fig. 1) or our mapping (fig. 7). indeed, with the few exposures available, it is extremely difficult to reconstruct the potential subcrop of permian or triassic strata, although in the north of wollaston forland, the basement is directly overlain by jurassic and cretaceous strata. to what extent thin permian, or triassic, deposits could have been eroded from uplifted regions, as those we now see, are, of course, unknown. indeed, haller (1971, fig. 145) appears to suggest that permian strata blanketed the whole region prior to rifting and erosion during the triassic. the discovery of reworked triassic palynomorphs in the jurassic succession at jurakløft, described above, provides evidence for the erosion suggested by haller (1971) and constrains its timing as late triassic to early jurassic. when considered in relation to the adjacent offshore basins, the northward continuation of the permian carbonates is consistent with what has been documented on the mid-norwegian continental shelf (bugge et al. 2002; müller et al. 2005) and predicted to exist on the east greenland shelf (rowan & jarvie 2020). previously, the most northerly triassic strata in east greenland were mapped as occurring on clavering ø, a short distance to the south of the sections described here, where deep marine clastics of the wordie creek formation overlie the permian succession. the wordie creek formation dominates the northern extremities of the triassic outcrop with the northernmost continental triassic occurring on laplace bjerg, geographical society ø, 73°n (andrews et al. 2019, 2020a). therefore, the presence of continental triassic (pingo dal group) on wollaston forland (74°22´n) significantly expands the extent of known continental facies of this age and suggests similar facies might be found in adjacent offshore basins. it should be remembered that the assignation of these facies to the triassic is largely on the basis of their lithology, and the difficulty in placing such a continental succession within the late permian or jurassic stratigraphy. therefore, although an early triassic age, comparable to the pingo dal group elsewhere, would seem most likely, a later triassic age cannot be ruled out. the presence of reworked triassic (carnian) palynomorphs within the jurassic certainly suggests that triassic cover was more extensive than the single outcrop identified here. this would be consistent with the more complete triassic successions recorded on the east greenland shelf (rowan & jarvie 2020) and the conjugate mid-norwegian shelf (müller et al. 2005). resolving the controls on the uplift and erosion is difficult on the data available. it might be considered that late triassic – early jurassic erosion in the region was the result of rift flank uplift and was, therefore, localised around the basin margins. however, an apparently significant change in fault orientation occurs between the triassic (ne–sw) and jurassic (nnw–sse) in east greenland (surlyk 2003; https://doi.org/10.34194/geusb.v47.6523 www.geusbulletin.org� andrews et al. 2021: geus bulletin 47. 6523. https://doi.org/10.34194/geusb.v47.6523 13 of 14 www.geusbul let in.org guarnieri et al. 2017; andrews et al. 2020b), and therefore, some form of structural re-organisation might be inferred, which could have resulted in uplift and erosion. a change in basin configuration is also recognised at this time in the barents sea and is attributed to the development of the novaya zemlya fold and thrust belt (müller at al. 2019). although it would be rather far afield for these events to have affected east greenland, it does demonstrate that significant tectonic events were occurring at this time in the region, and therefore, related adjustments in stress fields could have contributed to the development of the observed hiatus. the permpasset horst, which provides the exposure of the units described above, is defined in the east by the permpasset fault, and to the west by a ne–sw-aligned fault. these faults converge to the north. a further fault aligned e–w cuts the permpasset horst but displays relatively limited displacement. the ne–sw fault and the e–w fault post-date the jurassic and triassic, respectively, and the former is sealed by the albian fosdalen formation. their formation is, therefore, most likely related to early cretaceous rifting. following the rifting, a hiatus surface was developed and was subsequently draped with mid-albian strata of the fosdalen formation, during the post-rift thermal subsidence. movement on the permpasset fault can be demonstrated to postdate eocene basalt emplacement, but this does not preclude earlier episodes of movement. of further interest is the presence of sandstone clasts in the pebbly conglomerate at the base of the jurassic section as these may suggest that the permian and triassic sediments were exposed and being eroded during the jurassic and, therefore, imply that uplift of the region had already taken place in some form. glendonites are found along the basal hiatus surface, which marks the base of the cretaceous succession. this is the first record of glendonites in the cretaceous of east greenland, and these likely relate to the late aptian – early albian cooling event that is recorded throughout the arctic (rogov et al. 2017). this would suggest that the glendonites were, in fact, derived from older cretaceous strata during the generation of the hiatus surface. finally, the presence of permian and triassic clastic strata, in the crestal region of a major horst structure, that are, in turn, draped with cretaceous siltstones– mudstones, provides an important analogue for offshore hydrocarbon traps, expanding on the models presented by surlyk & korstgård (2013). acknowledgments this work was undertaken during a geus and mierl joint co-ordinated wollaston forland regional mapping project. ashton embry and tore grane klausen are thanked for their constructive reviews. funding statement geus and mierl joint funded the project. competing interests the authors declare no competing interests. author contributions sda: field data collection, writing and conceptualisation. hnh: field data collection, palynological analysis and manuscript editing. pg: field data collection, mapping (field and photo) and writing (contribution). kd: palynological analysis. sl: palynological analysis. pa: manuscript editing. references andrews, s.d. & decou, a. 2019: the triassic of traill ø and geographical society ø, east greenland: implications for north atlantic palaeogeography. geological journal 54(4), 2124–2144. https://doi.org/10.1002/ gj.3287 andrews, s.d., decou, a., braham, b., kelly, s.r.a., robinson, p., morton, a., marshall, j.e.a. & hyden, f. 2020a: exhumed hydrocarbon traps on the north atlantic margin: stratigraphy, palaeontology, provenance and bitumen distribution, an integrated approach. basin research 32(5), 1213–1233. https://doi.org/10.1111/bre.12424. andrews, s.d., kelly, s.r., braham, w. & kaye, m. 2014: climatic and eustatic controls on the development of a late triassic source rock in the jameson land basin, east greenland. journal of the geological society 171(5), 609–619. https://doi.org/10.1144/jgs2013-075 andrews, s.d., morton, a.c. & decou, a. 2020b: mid to late triassic evolution of the jameson land basin, east greenland. geological magazine. 158(5), 930–949 https://doi.org/10.1017/s001675682000093x bjerager, m., alsen, p., bojesen-koefoed, j., fyhn, m.b., hovikoski, j., ineson, j., nøhr-hansen, h., nielsen, l.h., piasecki, s. & vosgerau, h. 2020: cretaceous lithostratigraphy of north-east greenland. bulletin of the geological society of denmark 68, 37–93. https://doi.org/10.37570/ bgsd-2020-68-04 bugge, t., ringas, j.e., leith, d.a., mangerud, g., weiss, h.m. & leith, t.l. 2002: upper permian as a new play model on the mid-norwegian continental shelf: investigated by shallow stratigraphic drilling. aapg bulletin 86(1), 107–127. https://doi. 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(ed.): future petroleum provinces of the world. aapg memoirs 40, 629–659. https://doi.org/10.1306/ad461791-16f7-11d7-8645000102c1865d surlyk, f. & korstgård, j. 2013: crestal unconformities on an exposed jurassic tilted fault block, wollaston forland, east greenland as an analogue for buried hydrocarbon traps. marine and petroleum geology 44, 82–95. https://doi.org/10.1016/j.marpetgeo.2013.03.009 vigran, j.o., mangerud, g., mørk, a., worsley, d. & hochuli, p.a. 2014: palynology and geology of the triassic succession of svalbard and the barents sea. geological survey of norway special publication 14, 269 pp. vosgerau, h., bojesen-koefoed, j.a., petersen, h.i. & surlyk, f. 2000: forest fires, climate, and sea-level changes in a coastal plain-shallow marine succession (early–middle oxfordian jakobsstigen formation, northeast greenland). journal of sedimentary research 70(2), 408–418. https://doi.org/10.1306/2dc40919-0e47-11d7-8643000102c1865d https://doi.org/10.34194/geusb.v47.6523 www.geusbulletin.org� https://doi.org/10.1016/s0928-8937(01)80008-4 https://doi.org/10.1016/s0928-8937(01)80008-4 https://doi.org/10.1144/jgs2015-028 https://doi.org/10.1016/j.tecto.2019.04.006 https://doi.org/10.1016/s0928-8937(05)80048-7 https://doi.org/10.1016/s0928-8937(05)80048-7 https://doi.org/10.34194/bullggu.v166.6722 https://doi.org/10.34194/bullggu.v166.6722 https://doi.org/10.1017/s0016756819001043 https://doi.org/10.1016/j.revpalbo.2015.05.001 https://doi.org/10.1017/s0016756819000906 https://doi.org/10.1017/s0016756819000906 https://doi.org/10.34194/geusb.v5.4808 https://doi.org/10.34194/seriec.v10.7104 https://doi.org/10.34194/seriec.v10.7104 https://doi.org/10.1016/j.cretres.2016.11.011 https://doi.org/10.1016/j.marpetgeo.2020.104339 https://doi.org/10.1016/j.marpetgeo.2020.104339 https://doi.org/10.1144/0016-764903-063 https://doi.org/10.1144/0016-764903-063 https://doi.org/10.34194/geusb.v41.4353 https://doi.org/10.34194/geusb.v41.4343 https://doi.org/10.34194/geusb.v41.4343 https://doi.org/10.1111/j.1365-3091.2001.00352.x https://doi.org/10.1144/gsl.sp.1990.055.01.05 https://doi.org/10.1144/gsl.sp.1990.055.01.05 https://doi.org/10.34194/geusb.v1.4674 https://doi.org/10.34194/geusb.v1.4674 https://doi.org/10.1306/ad461791-16f7-11d7-8645000102c1865d https://doi.org/10.1016/j.marpetgeo.2013.03.009 https://doi.org/10.1306/2dc40919-0e47-11d7-8643000102c1865d the permian to cretaceous succession at permpasset, wollaston forland: the northernmost permian and triassic in north–east greenland abstract 1 introduction and geological background 2 methodology 3 sedimentology and stratigraphy 3.1 permpasset horst 3.2 north west of the permpasset horst 4 structure of the permpasset horst 5 discussion and conclusions acknowledgments references figures fig. 1 geological map of wollaston forland (modified after escher 2001). the black-lined box marks the position of the map illustrated in fig. 7. the inset map indicates the position of wollaston forland (wf) on the east greenland coast as well as other regions mentioned in the text: cø: clavering ø. hwh: hold with hope. gsø: geographical society ø. tl: traill ø. sl: scoresby land. jl: jameson land. fig. 2 overview of the permpasset outcrop and photographs illustrating the key lithologies of the permian succession. a: view west across permpasset illustrating the stratigraphic and structural relationships and the position of the logged sections provided in figs 3 and 6. p: permian. tr: triassic. jr: jurassic. k: cretaceous. b: brecciated limestone and gypsum, including progressively larger blocks upwards (karstryggen formation). c: bedded selenite overlying the first brecciated interval (karstryggen formation). d: bedded limestones recording a deepening and freshening of marine conditions, perhaps forming a correlative of the wegener halvø formation. fig. 3 logged section from the central outcrops of the permpasset exposures. the location of the section is indicated in fig. 2a. sample numbers are included adjacent to the log. fig. 4 photographs illustrating the key lithologies of the uppermost permian, and the triassic, jurassic and cretaceous stratigraphy. a: interbedded shelly limestones and sandstones forming the top of the permian succession (schuchert dal formation). b: buff-coloured planar and trough cross-bedded sandstones with intraclast lags commonly present at bed bases (triassic pingo dal group?). c: pebbly sandstone with abundant rounded quartz clasts and occasional sandstone clasts forming the base of the jurassic (jakobsstigen formation). d and e: different glendonite forms from the base of the cretaceous succession. f: graded lamination and contorted intraclasts within the siltstones of the fosdalen formation. g: current ripples with possible signs of storm reworking in sandstone beds weathering out of the fosdalen formation. h: flute marks and thalassinoides ichnofossils preserved on bed bases in the fosdalen formation. fig. 5 examples of reworked triassic pollen. taxon name is followed by sample and slide number and england finder coordinates. the scale bar is 20 μm. a: aulisporites astigmosus, 589748:4, u58/3. b: protodiploxypinus gracilis, 589757:5, h18/1. c: lunatisporites sp., 589754:5, h53/2. d: protodiploxypinus sp. cf. p. gracilis, 589757:4, s43/1. fig. 6 logged sections from the northern portion of permpasset. the location of permpasset w is provided in fig. 2a. pal/eo: paleocene or eocene. sample numbers are included adjacent to the log. fig. 7 geological map of the permpasset region and associated cross-section. a: detailed geological map of permpasset from fieldwork and 3d–photogeology. all heights given are metres a.s.l. caledonian migmatites are unconformably overlain by upper permian limestones of the karstryggen formation followed by a sandy succession of the schuchert dal formation (not possible to distinguish in the northern portion of the outcrop due to erosion) and red beds of the pingo dal group. mid–albian mudstones of the fosdalen formation unconformably cover both the pingo dal group and jakobsstigen formation and post-date the ne–sw-trending normal fault. the n–s-trending permpasset fault was active in palaeogene times as shown by almost 400 m offset of platea basalts to the east and responsible for the westward tilting of the successions. b: wsw–ene geological cross-section of permpasset showing structural relationships and the geometry of the fault-bounded permpasset horst. monograph japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 1 of 154 episodic burial and exhumation in north-east greenland before and after opening of the north-east atlantic peter japsen*1 , paul f. green2 , johan m. bonow3,4 , morten bjerager1 , john r. hopper1 1geological survey of denmark and greenland (geus), copenhagen, denmark. 2geotrack international, brunswick west, victoria, australia. 3geovisiona ab, bro, sweden. 4uppsala university, uppsala, sweden. abstract the geology of north-east greenland (70–78°n) exposes unique evidence of the basin development between the devonian collapse of the caledonian orogen and the extrusion of volcanics at the paleocene–eocene transition during break-up of the north-east atlantic. here we pay special attention to unconformities in the stratigraphic record – do they represent periods of stability and non-deposition or periods of subsidence and accumulation of rocks followed by episodes of uplift and erosion? to answer that and other questions, we used apatite fission-track analysis and vitrinite reflectance data together with stratigraphic landscape analysis and observations from the stratigraphic record to study the thermo-tectonic history of north-east greenland. our analysis reveals eight regional stages of post-caledonian development: (1) late carboniferous uplift and erosion led to formation of a sub-permian peneplain covered by coarse siliciclastic deposits. (2) middle triassic exhumation led to removal of a thick cover including a considerable thickness of upper carboniferous – middle triassic rocks and produced thick siliciclastic deposits in the rift system. (3) denudation at the transition between the early and middle jurassic affected most of the study area outside the jameson land basin and produced a weathered surface above which middle–upper jurassic sediments accumulated. (4) earliest cretaceous uplift and erosion along the rifted margin and further inland accompanied the mesozoic rift climax and produced coarse-grained sedimentary infill of the rift basins. (5) mid-cretaceous uplift and erosion initiated removal of cretaceous post-rift sediments that had accumulated above the mesozoic rifts and their hinterland, leading to cooling of mesozoic sediments from maximum palaeotemperatures. (6) end-eocene uplift was accompanied by faulting and intrusion of magmatic bodies and resulted in extensive mass wasting on the east greenland shelf. this event initiated the removal of a thick post-rift succession that had accumulated after break-up and produced a peneplain near sea level, the upper planation surface. (7) late miocene uplift and erosion, evidenced by massive progradation on the shelf, resulted in the formation of the lower planation surface by incision below the uplifted upper planation surface. (8) early pliocene uplift raised the upper and the lower planation surfaces to their present elevations of about 2 and 1 km above sea level, respectively, and initiated the formation of the present-day landscape through fluvial and glacial erosion. additional cooling episodes of more local extent, related to igneous activity in the early eocene and in the early *correspondence: pj@geus.dk received: 28 nov 2018 accepted: 26 mar 2021 published: 8 oct 2021 keywords: erosion, hiatus, peneplain, uplift, tectonics abbreviations: a.s.l.: above sea level afta: apatite fission-track analysis arco: atlantic richfield company epcm: elevated passive continental margin es: etch surface imu: intra-miocene unconformity lps: lower planation surface odp: ocean drilling program pdmf: post-devonian main fault pse: projected surface elevation ups: upper planation surface vr: vitrinite reflectance vt: vertical transect geus bulletin is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: catherine n. jex (geus, denmark) reviewed by: finn surlyk (university of copenhagen, denmark), sierd cloetingh (utrecht university, netherlands) funding: see page 85 competing interests: none declared additional files: see page 85 https://orcid.org/0000-0001-5478-534x https://orcid.org/0000-0003-0547-0565 https://orcid.org/0000-0003-3180-8857 https://orcid.org/0000-0003-3188-7583 https://orcid.org/0000-0003-1689-7820 japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 2 of 154 www.geusbul let in.org miocene, primarily affected parts of northern jameson land. the three earliest episodes had a profound impact beyond greenland and accompanied the fragmentation of pangaea. younger episodes were controlled by plate-tectonic processes, possibly including dynamic support from the iceland plume. our results emphasise that gaps in the stratigraphic record often reflect episodes of kilometre-scale vertical movements that may result from both lithospheric and sub-lithospheric processes. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 3 of 154 www.geusbul let in.org 1. introduction north-east greenland (70–78°n) exposes unique evidence of the basin development between the devonian collapse of the caledonian orogen and the extrusion of volcanics at the paleocene–eocene transition during break-up of the north-east atlantic (figs 1, 2; haller 1971; escher & watt 1976; surlyk 1990, 2003; stemmerik 2000; higgins et al. 2008; larsen et al. 2008a; henriksen et al. 2009). the region hosts spectacular mountains reaching 2.8 km above sea level (a.s.l.) within the study area, and the highest mountain in greenland, gunnbjørn fjeld (3.7 km a.s.l., c. 68°n), carved in palaeogene volcanics, is located just south of the region. of particular interest are the development of the modern topography along this passive continental margin and its importance for the formation of the 3 km thick greenland ice sheet. the stratigraphic record of north-east greenland includes several major unconformities. for example, there are no pre-quaternary sediments younger than the paleocene – lowermost eocene rocks within our study area, leaving a gap of about 50 myr in the stratigraphic record (nøhr-hansen et al. 2011; larsen et al. 2014; stoker et al. 2017). a key question in regard to the evolution of the region is whether such gaps represent periods of stability and non-deposition or periods in fig. 1 location of study area in north-east greenland (70–78°n) and present-day topography of the bedrock of greenland. the burial and exhumation history of various parts of greenland have been investigated using apatite fission-track analysis (afta) data in a series of publications, as indicated on the map (japsen et al. 2005, 2006, 2009, 2010, 2014, 2021 (this volume); green et al. 2011, 2013, 2014). the load of the greenland ice sheet causes up to 850 m subsidence of the bedrock topography of central greenland. peripheral bulging caused by this ice loading has a negligible effect (<20 m) on the elevation of the greenland margins (medvedev et al. 2013). elevation data from amante & eakins (2009). 70°n 65°n 60°n 75°n this study japsen et al. 2021 (this volume) japsen et al. 2014 japsen et al. 2005, 2006, 2009, 2010 green et al. 2011, 2013 80°n elevation (m) 20°w 30°w 40°w 50°w 60°w –3000 – –1000 – –2000 – 0 – –4000 – 1000 – 2000 – 3000 – green et al. 2014 fig. 1-1 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 4 of 154 www.geusbul let in.org which significant thicknesses of rock were deposited and then removed as a result of uplift and erosion? in either case, such intervals represent gaps in our understanding of the full geological development of the region. when the gaps represent burial followed by exhumation, we are left with a censored view of the geological history, and of the tectonic processes that shaped the surface of the earth as we know it today. here we will investigate the nature of major post-devonian unconformities in the stratigraphic record of north-east greenland. north-east greenland is a typical example of an elevated passive continental margin (epcm), which occur in different climatic zones and along margins of different ages across the world, including scandinavia, brazil and south-east australia (japsen et al. 2012a; green et al. 2013, 2018). these margins typically have mesozoic–palaeogene rift systems parallel to the coast (both onshore and offshore in the case of north-east greenland; hamann et al. 2005; døssing et al. 2016) and a transition from continental to oceanic crust farther offshore. also in north-east greenland, escarpments separate low-lying coastal areas from the mountains that are characterised by elevated plateaus and deeply incised valleys (bonow & japsen 2021, this volume). furthermore, synand post-rift sediments at the landward margin of these rifts dip away from the epcm towards the rifts and are truncated by one or more unconformities (japsen & chalmers 2000). the timing and nature of the processes driving the formation of epcms have been topics of much debate in recent years (nielsen et al. 2009, 2010; chalmers et al. 2010; pedersen et al. 2012; japsen et al. 2012a, 2013, 2019; green et al. 2013, 2018; colli et al. 2014, 2016; braun 2018, 2019). are they remnants of ancient mountain chains? are these mountains remnants of rift shoulders from the time of rifting and break-up? or are they young expressions of dynamic topography or of far-field transmission of stress (braun 2010; cloetingh & burov 2010; colli et al. 2016)? can elevated topography be sustained on timescales of hundreds or only tens of millions of years? in this study, we examine the formation of the coastal mountains in north-east greenland. we present new apatite fission-track analysis (afta®) data that shed light on the episodes of subsidence, burial, uplift and erosion that affected the region both prior to and after break-up of the north-east atlantic. we integrate stratigraphic landscape analysis (lidmar-bergström et al. 2013) with new and published afta and vitrinite reflectance (vr) data. these approaches provide complementary information (green et al. 2013). while stratigraphic landscape analysis provides a relative denudation chronology, afta can provide absolute timing constraints for the onset of denudation episodes, which initiate the formation of peneplains identified from the landscape analysis. the age for the final formation of the peneplains can be constrained by cover rocks, where present. we have previously published results from southern east greenland (68–71°n) based on integration of afta data and stratigraphic landscape analysis (bonow et al. 2014; japsen et al. 2014) and results from south-east greenland (61– 66°n) based on afta data alone (green et al. 2014). in part one of this volume, we presented an outline of the thermo-tectonic development of the wandel sea basin, north greenland (japsen et al. 2021, this volume). we also consider the impact of exhumation and igneous activity on the hydrocarbon systems in the sedimentary basins of north-east greenland. the outcropping sedimentary basins of north-east greenland have long been recognised as potential analogues for understanding offshore basins and their hydrocarbon systems; particularly relative to those on the conjugate norwegian margin (surlyk 1990; stemmerik & worsley 2005). the us geological survey concluded that the north-east greenland shelf could be an important future petroleum province but stressed that cenozoic uplift and erosion are important risk factors for the development of the hydrocarbon system (gautier et al. 2011). christiansen (2011) found uplift to be the main exploration risk in east and north-east greenland, but he also assigned a significant uncertainty to the limited stratigraphic control due to lack of geophysical and well data. major oil companies acquired several licences for hydrocarbon exploration off north-east greenland between 76 and 80°n in 2012 and 2013 (bojesen-koefoed et al. 2014). however, this acreage has now been relinquished. finally, we consider the impact of the reconstructed history of uplift and subsidence on the development of the greenland ice sheet. it is unclear whether the ice sheet persisted for millions of years or whether greenland has been ice-free for extended periods during the pleistocene (bierman et al. 2016; schaefer et al. 2016; christ et al. 2021). numerical models of the development of the greenland ice sheet traditionally use the present-day bedrock geometry as a boundary condition (letreguilly et al. 1991). but recently, solgaard et al. (2013) investigated the influence of late neogene mountain building in greenland on the initiation of the ice sheet. we review these results. the stratigraphic record and the large-scale landscapes of the study area form the backbone of our investigation (chapters 2 and 3; bonow & japsen 2021, this volume). a synthesis of afta data and associated thermal-history interpretations from 217 outcrop samples is presented in chapter 4. we integrate these results with observations from the geological record and with results from the stratigraphic landscape analysis (chapters 5 and 6) similar to our approach in studies http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 5 of 154 www.geusbul let in.org of central west greenland and southern east greenland (japsen et al. 2006, 2014). in chapters 7 and 8, we discuss the mesozoic burial and exhumation history and the effects of exhumation on the hydrocarbon prospectivity. in chapter 9, we compare our results from northeast greenland with those from west greenland and from southern scandinavia and other areas. lastly, in chapter 10, we discuss several general issues raised by the results presented here including the development of the greenland ice sheet. liverp o o l lan d gåse�ord 100 km gåseland kap stosch loch fyne sc hu ch er t d al payer land giesecke bjerge dombjerg geographical society ø gauss halvø wegener halvø wager plateau kangerlussuaq mestersvig norsketinden 2.8 km 25°w30°w 69°n 72°n 75°n 69°n 72°n 75°n 20°w 25°w35°w 30°w 20°w germania land dove bugt store koldewey shannon kuhn ø hochstetter forland wollaston forland clavering ø hold with hope myggbukta hudson land ymer ø ella ø jameson land renland lyell land nathorst land th. sørensen land milne land geikie plateau kap dalton kap brewster gunnbjørn fjeld 3.7 km b l o s s e v i l l e k y s t s c o r e s b y s u n d ø�ord nordvest�ord k o n g o s c a r f j o r d traill ø arnold escher land stauning alper sabine ø kejser franz joseph fjord 0–300 elevation (m a.s.l.) 33–1500 1500+ japetus bjerg werner bjerge fig. 2 elevation and place names of the study area in north-east greenland (north of scoresby sund) and of southern east greenland (between scoresby sund and kangerlussuaq). http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 6 of 154 www.geusbul let in.org 2. post-caledonian geology the roots of the east greenland caledonian fold belt are well exposed between 70 and 81°30´n as a 1300 km long and up to 300 km wide belt parallel to the coast (higgins et al. 2008). large regions of the fold belt are characterised by reworked precambrian basement rocks overlain by proterozoic and lower palaeozoic sedimentary rocks, all of which form parts of westward-directed, major thrust sheets formed during the climax of collision between laurentia and baltica in mid-silurian times (the caledonian orogeny). extensional collapse of an overthickened crust initiated basin formation in east greenland. from middle devonian times, more than 8 km of mainly coarse siliciclastic sediments accumulated in the continental old red sandstone basin in north-east greenland (larsen et al. 2008a). the initial phase of the collapse resulted in major exhumation, evident from the deposition of old red sandstone in the eifelian (early middle devonian; 393–388 ma) shortly after a phase of high-pressure metamorphism at c. 405 ma (gilotti & mcclelland 2008). in this chapter, we briefly describe the post-caledonian geology of north-east greenland. 2.1 carboniferous–cretaceous cover and unconformities a series of carboniferous–mesozoic sedimentary basins developed in east greenland following initial, post-caledonian devonian deposition (fig. 3). the basin evolution was governed by a system of old first-order faults, some of which were formed by reactivation of caledonian or older thrusts and weakness zones (surlyk 2003). new faults formed during continued extensional movements. in the south, a very deep basin was formed in the jameson land area in late devonian – early permian times (see the stratigraphic scheme in fig. 4). this area then acted as a wide, coherent platform in late permian – jurassic times. subsidence was accommodated mainly along basin margin faults, and there was no syn-sedimentary faulting of the platform. this is in contrast to the region north of jameson land, notably the wollaston forland – kuhn ø area, which was fragmented into narrow tilted blocks during middle jurassic – earliest cretaceous rifting. tectonic models of the carboniferous–permian basins of north-east greenland have emphasised rifting and half graben formation along the so-called post-devonian main fault (pdmf; vischer 1943) – a north to north-easttrending fault, which appears as a marked scarp in the terrain, and defines the present-day western boundary of the sedimentary basins north of kong oscar fjord. south of the fjord, it was termed the stauning alper fault. we will refer to these two faults as the post-devonian main fault system. however, bütler (1955) observed offsets of tertiary dykes on ymer ø and on geographical society ø along the post-devonian main fault and concluded that the main faulting took place during and after the emplacement of the tertiary dykes. bütler (1955, p. 128) wrote that “the term ‘post-devonian main fault’ for this part of the large fault line seems therefore rather indefinite”. during rifting in the latest devonian and carboniferous, deposition took place in a system of north-south elongated and relatively narrow half-grabens. the carboniferous redbeds of the traill ø group in north-east greenland accumulated during initial rifting of the northern north atlantic (vigran et al. 1999; stemmerik 2000). deposition took place in a system of westward tilted halfgrabens, exposed from clavering ø in the north to jameson land in the south (surlyk et al. 1984; stemmerik et al. 1990, 1991). a hiatus spans the mid-carboniferous in east greenland (stemmerik et al. 1991), and upper carboniferous deposits rest on basement on germania land and holm land farther north (piasecki et al. 1994). a major hiatus spanning the latest carboniferous to early permian separates tilted strata of the traill ø group from the overlying upper permian conglomerates of the huledal formation (surlyk 1990; vigran et al. 1999; stemmerik 2000). this unconformity marks the most profound change in tectonic style and overall depositional environment in the post-caledonian development of east greenland (surlyk 1990). the succession below the unconformity was uplifted and deeply eroded along the basin margins during a final phase of rotational block-faulting, and a vast peneplain developed (haller 1971; larsen 1988; surlyk 1990). the widespread deposits of the huledal formation accumulated on the peneplain with pronounced unconformable contact and marks the transition from a long period of crustal extension to a period of subsidence governed mainly by thermal relaxation of the rifted crust (surlyk 1990). the mid–late permian røde ø conglomerate downfaulted against the basement of the southern interior (west of milne land), is an equivalent to the huledal formation, and thus testifies to the former regional extent of these deposits (stemmerik & piasecki 2004). in the late permian, the sea transgressed the area for the first time since the caledonian orogeny, allowing deposition of the foldvik creek group. haller (1971, p. 320) summarised this change in the structural development and emphasised that the caledonian mountain had disappeared by then: “when the upper permian sea transgressed across the present outer fjord region most http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 7 of 154 www.geusbul let in.org of the intra-permian fault escarpments had been worn away. the late palaeozoic block mountain ranges had disappeared and given place to lowlands. in the course of this permian denudation considerable quantities of carbono-permian molasse deposits must have been removed, in particular from the protruding edges of tilted fault blocks”. the marine upper permian – lower triassic succession in east greenland is up to 1 km thick and includes an upper permian carbonate and mudstone-dominated succession and a lower triassic sandstone and mudstone succession (e.g. stemmerik et al. 1993; seidler et al. 2004; bjerager et al. 2006). marine organic-rich mudstones of the upper permian ravnefjeld formation of the foldvik creek group are widespread in east greenland. the mudstones were deposited along the axis of a more than 400 km long and at least 80 km wide, marine basin, flanked by shallow-water carbonate platforms. rifting began at the permian–triassic transition (surlyk et al. 2017). two early triassic marine rifting events were recognised in east greenland by seidler et al. (2004), who correlated the later event with rhyolitic intrusions in liverpool land with an ar-ar age of 244 ± 4 ma. early triassic rifting was followed by middle triassic deposition of the coarse-grained, continental conglomerates and sandstones of the pingo dal group, which overlies the early triassic marine mudstones of the wordie pre-devonian rocks outlier fault devonian sediments carboniferous sediments cretaceous sediments jurassic sediments triassic sediments permian sediments palaeogene basalts palaeogene intrusives cenozoic sediments quaternary sediments sea ice 35°w 30°w 25°w 20°w 20°w25°w30°w 15°w 76 °n 74 °n 72 °n 70 °n 76 °n 74 °n 72 °n 70 °n 50 km liverpool land kap brewster d f pd m f pd m f sa f jameson land schuchert dal wegener halvø malmbjerget myggbukta payer land traill ø gauss halvø kong oscar fjord kejser franz joseph fjord geographical society ø hold with hope clavering ø kuhn ø hvalros ø sabine ø shannon dombjerg hochstetter forland dove bugt store koldewey germania land wollaston forland scoresby sund milne land renland stauning alper hudson land arnold escher land gåseland ymer ø fig. 3 geological map of north-east green land (escher & pulvertaft 1995; henriksen 2003; surlyk 2003; christoffersen & jepsen 2007). major post-devonian faults are shown (surlyk 2003; guarnieri 2015). a middle jurassic outlier on germania land is located just north of the map frame (bojesen-koefoed et al. 2012). df: dombjerg fault. pdmf: post-devonian main fault. saf: stauning alper fault. we refer to the two latter faults collectively as the post-devonian main fault system. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 8 of 154 www.geusbul let in.org ma era stage/ age 0 10 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 210 220 230 240 250 260 270 280 290 300 310 320 330 340 350 paleocene eocene oligocene miocene pliocene pleistocene series/ epoch system/ period middle middle upper upper lower upper guadalupian lopingian lower lower middle mississippian upper mississippian lower penns. middle penns. upper penns. cisuralian lower mississippian messinian tortonian serravallian langhian aquitanian chattian rupelian priabonian bartonian lutetian thanetian selandian danian maastrichtian campanian santonian coniacian cenomanian albian aptian hauterivian valanginian berriasian tithonian kimmeridgian callovian bathonian bajocian aalenian pliensbachian sinemurian hettangian car nian ladinian anisian olenekian induan rhaetian norian roadian wordian capitanian wuchiapingian changhsingian kungurian artinskian sakmarian asselian gzhelian kasimovian moscovian bashkirian serpukhovian visean tournaisian oxfordian barremian toarcian turonian ypresian burdigalian pa le o zo ic m es o zo ic c en o zo ic c a r bo n if er o u s pe r m ia n tr ia ss ic ju r a ss ic c r et a c eo u s pa la eo g en e n eo g en e wollaston foreland – kuhn ø w e s n ? ? wollaston forland gp vardekløft gp bernbjerg fm foldvik creek gp store koldewey w e pelion fm traill ø – hold with hope s n w e traill ø gp foldvik creek huledal fm gp gråklint mb pingo dal gp gipsdalen gp fleming fjord gp vardekløft gp bernbjerg fm wollaston forland gp steensby bjerg fm home forland gp brorson halvø gp jackson ø gp månedal fm pelion fm jameson land basin s n ? ? w e kap stewart gp neill klinter gp traill ø gp foldvik creek wordie creek gp gp vardekløft gp pelion fm hesteelv fm raukelv fm hareelv fm sw ne milne land hartz fjeld fm kap leslie fm charcot bugt fm fig. 4 caption and legend on the next page. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 9 of 154 www.geusbul let in.org creek group above a pronounced unconformity, for example in northern jameson land (clemmensen et al. 2020). the deposition of the pingo dal group was the result of a major phase of basin-margin uplift and rapid fault-controlled basin subsidence in early triassic time (late scythian) leading to the formation of a north– south-trending intermontane graben, about 300 km long and 75 km wide (clemmensen 1980; surlyk 1990). the transgressive development during the deposition of the rhaetian – lower bajocian pre-rift succession was terminated by erosion that took place broadly at the early–middle jurassic boundary (surlyk 2003; surlyk et al. 2021). the uplift event that affected the area north of jameson land at this time is reflected in a short, mid-bajocian hiatus between the sortehat and pelion formations. on kuhn ø, the marine pelion formation rests on peneplained basement (surlyk 2003). the subsequent onset of a long-term rifting episode began in the middle jurassic, peaked in the late jurassic and persisted into the earliest cretaceous (surlyk 2003). the southern part of the basin complex in jameson land contains a relatively complete rhaetian–ryazanian succession deposited during uniform subsidence, only affected by minor tilting during middle jurassic – earliest cretaceous rifting. in contrast, rhaetian – lower jurassic deposits are absent on milne land and north of jameson land, and the northern region was fragmented into strongly tilted fault blocks during the protracted rift event. on milne land and in areas north of jameson land, middle and upper jurassic sandstones (e.g. the charcot bugt and pelion formations) onlap triassic and upper permian sediments and crystalline basement. surlyk (1977, 1978a, 2003) speculated whether the rhaetian – lower bajocian succession may have been removed by erosion during late early jurassic domal uplift of northern east greenland followed by subsidence and middle jurassic onlap. the main sources for the jurassic sedimentary basin complex in north-east greenland were precambrian and caledonian metamorphic and intrusive rocks and thick proterozoic sedimentary successions (surlyk 2003). source terranes of minor importance were devonian conglomerates, sandstones and mudstones as well fig. 4 (continued) lithostratigraphic scheme for north-east greenland. the columns for the jameson land basin, traill ø – hold with hope, wollaston forland – kuhn ø are based on surlyk (1990) and on later compilations (gautier et al. 2011; hopper et al. 2014; stoker et al. 2017; bjerager et al. 2020; clemmensen et al. 2020; surlyk et al. 2021). intrusive magmatism after hopper et al. (2014) and larsen et al. (2014). tectonic and orogenic events after hopper et al. (2014). base of pingo dal group placed at 245 ma (seidler et al. 2004; oftedal et al. 2005); gråklint beds at 232 ma, mid-carnian (andrews et al. 2014). timescale after gradstein et al. (2012). penns: pennsylvanian. base-of-slope fans channel fills marine mudstone marine carbonate, including chalk seas deep marine sandstone (beyond shelf break) shallow marine sandstone (shelf) arid terrestrial sedimentary environments lithological symbols other symbols anhydrite/gypsum coal layers sandstone siltstone mudstone/shale organic rich mudstone/shale limestone prograding shelf-margin sediment wedge volcanics extrusive igneous rocks mainly mafic rifting/extension humid terrestrial/ marginal marine dolomite/evaporites/ limestone conglomerate bioherm ice-rafted dropstone intrusive magmatism tectonic and orogenic events inversion/compression/ uplift http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 10 of 154 www.geusbul let in.org as carboniferous sandstones and mudstones. triassic conglomerates, sandstones and mudstones constituted an additional source in the early jurassic due to uplift of the areas north of jameson land. the most important phase of extensional tectonics since the carboniferous – early permian, culminated in middle volgian times (tithonian–berriasian transition around 145 ma; surlyk & ineson 2003), at the jurassic–cretaceous boundary (surlyk 1978b, 2003; surlyk & korstgård 2013). the area north of kong oscar fjord broke up into tilted blocks, which became progressively fragmented and narrower with time, and deposition was almost exclusively coarse grained. the jameson land area, south of kong oscar fjord, underwent only mild tilting but enormous quantities of coarse sands were shed into the basin in the tectonically active period from middle volgian to late ryazanian time (corresponding to late tithonian – late berriasian). the tectonic phase faded out in the latest valanginian. post-valanginian, cretaceous sediments crop out from traill ø in the south to store koldewey in the north. they comprise a mudstone-dominated succession up to several kilometres thick with significant coarse-grained units intercalated (nøhr-hansen 1993; stemmerik et al. 1993; bjerager et al. 2020). post-valanginian sedimentation was controlled by transgression and thermal contraction, and the main sedimentation pattern reflects the formation of a broad thermal sag above the volgian–valanginian fault blocks (surlyk 1990). prominent coarse-grained units, for example the rold bjerge and månedal formations, were suggested to represent significant cretaceous, possibly rift-associated, fault pulses superimposed on the long-term regional subsidence (surlyk & noe-nygaard 2001). recently, the cretaceous succession has been reviewed biostratigraphically (nøhr-hansen et al. 2019) and subdivided into three main tectonostratigraphic units represented by the upper hauterivian – middle albian brorson halvø group, the middle albian – coniacian home forland group and the turonian – lower maastrichtian jackson ø group (bjerager et al. 2020). the brorson halvø group is dominated by slope and basinal, mud-dominated deposits, filling the sea-floor relief inherited from the late jurassic – early cretaceous rift phase (bjerager et al. 2020). the steensby bjerg formation occurs locally on hold with hope and comprises a deltaic – shallow marine sand-dominated system that developed in an evolving relay ramp setting during the barremian – early albian (kelly et al. 1998; larsen et al. 2001). the fault-related mega-breccia of the re-dated aptian rold bjerge formation (surlyk & noe-nygaard 2001) occurs locally on traill ø. a pronounced mid-albian (c. 105 ma) unconformity marks the base of the overlying home forland group and is linked to localised fault re-activation and a drowning event. the group is dominated regionally by marine deep-water mud-dominated deposits of the middle albian – coniacian fosdalen formation, but locally fault associated conglomerates occur at the base on clavering ø (bjerager et al. 2020). the base of the jackson ø group rests on a prominent turonian erosional unconformity and the group represents a new phase of fault activity. on traill ø, the basal strata comprise submarine slope breccias and conglomerates of the månedal formation (surlyk & noe-nygaard 2001). in eastern hold with hope, the basal part is marked by the coniacian condensed nanok member (stemmerik et al. 1993, kelly et al. 1998). turbidite sandstones and basinal mudstones succeed these units. the lower campanian leitch bjerg formation of channelised slope-apron conglomerates occurs locally on eastern geographical society ø (bjerager et al. 2020). campanian–maastrichtian strata are also present in the kangerlussuaq area to the south (soper & costa 1976; larsen et al. 2001). 2.2 cenozoic cover and unconformities the final north-east atlantic rift episode was initiated near the campanian–maastrichtian boundary and lasted until continental separation at the paleocene–eocene transition, c. 56 ma (skogseid et al. 2000). more than 6 km of flood basalts (the main basalts) were extruded during rapid subsidence in the area along blosseville kyst, south of scoresby sund (larsen et al. 1989; pedersen et al. 1997; larsen & tegner 2006; brooks 2011). the first volcanics were partly erupted in a marine environment in paleocene time. furthermore, there are marine sediments within the youngest flood basalts, the lower eocene igtertivâ formation, and within the lower eocene to lower oligocene kap dalton group (soper & costa 1976; nielsen et al. 1981; larsen et al. 1989, 2002, 2013). the presence of marine deposits above and below the main basalts, which were extruded on a largely horizontal plain, implies that the landscape was low-lying during the volcanic eruptions. consequently, subsidence approximately kept pace with the outpouring lavas, so that the land surface at the end of the eruptions remained close to sea level (pedersen et al. 1997; brooks 1985, 2011; see review by bonow et al. 2014). a phase of middle miocene volcanism is documented by the presence of the lava flows of the vindtop formation (c. 14–13 ma; storey et al. 2004) that crop out on nunataks within a small area at 2.7–2.9 km a.s.l. (c. 69°n). remnants of the main basalts are also present in gåseland and in milne land. further north, lavas erupted at the paleocene–eocene transition and in the early eocene between 73 and 75°n (geographical society ø to shannon; fig. 5; larsen et al. 2014). outliers of palaeogene basalts are located on nunataks far from the coast (74°n, 28°w; fig. 3). http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 11 of 154 www.geusbul let in.org larsen et al. (1989) studied zeolites, the late-stage minerals that fill cavities in basalts, in the scoresby sund region, to infer amounts of removed basalts since their extrusion. zeolite isograds, which are contours that represent the first occurrence of a zeolite mineral, delineate regionally extensive mineral zones throughout the main basalts between scoresby sund and kangerlussuaq. these zones are essentially uniform in thickness and transgress lava stratigraphy (larsen et al. 1989; neuhoff et al. 1997). analysis of the isograds suggests that 400 m of the basalts were removed by erosion in the area just south of scoresby sund, whereas the erosion increases to 700 and 900 m towards the north, across gåseland and milne land in agreement with the gradual truncation of the basalt units in that direction (bonow et al. 2014). the zeolitisation thus had a regional character and took place before the tilting and erosion of the lava sequence (larsen et al. 1989). major offsets post-dating palaeogene basalts and intrusions occur along the pdmf and adjacent faults systems, for example more than 500 m on western hold with hope (bütler 1955; upton et al. 1980; hartz et al. 2006; surlyk & noe-nygaard 2001). scattered occurrences of palaeogene sediments are found in north-east greenland, where they unconformably overlie cretaceous sediments and are capped by palaeogene basalts. nøhr-hansen et al. (2011) showed that the marine, palaeogene sediments (now up to 700 m a.s.l.) on wollaston forland, hold with hope and sabine ø, comprise paleocene and lowermost eocene strata with a hiatus that probably spans the major part of the selandian and thanetian. nøhr-hansen et al. (2011) found indications of uplift and erosion at the maastrichtian–paleocene transition because the palaeogene sediments contained large amounts of reworked late maastrichtian marine palynomorphs. furthermore, they noted that the ages of the youngest palaeogene sediments and the oldest flood basalts are almost identical. stratigraphically extensive, ice-rafted debris, including macroscopic dropstones, occur in late eocene to early oligocene sediments from the norwegian–greenland sea, indicating sediment rafting by continental ice, rather than sea ice, and east greenland as the likely source (eldrett et al. 2007). however, continental ice on east greenland was probably restricted to alpine outlet glaciers across the eocene–oligocene transition (eldrett et al. 2009). material from offshore drill cores suggests that full glacial conditions were established in south-east 10 ma ma 20 30 40 50 60 20 30 40 50 60 63°n 64°n 65°n drift rift 66°n 67°n 68°n 69°n 70°n 71°n 72°n 73°n 74°n 75°n vindtop fm malmbjerget hvalros ø borgtinderne bjørn flammefjeld wiedemann fjord kangerlussuaq granite ejnar mikkelsen nephelinite diatreme ikâsangmit nûk myggbukta, late sheet werner bjerge kap simpson kap parry kap broer ruys (myggbukta, inferred) upper series ørsted dal mols bjerge shannonnunataks freycinet bjerg svinhufvud bjerge lower plateau lava series igtertivâ fm above cgl. igtertivâ fm below cgl. clasts in conglomerate pr. wales bj. fm gardiner granitekærven skaergaard sulugssut dykes imilik ii kap edvard holm nordre aputitêq igtutarajik pâtûlâjivit kruuse fjord kap gustav holm hole 988a hole seg01 hole 989b hole 990a lower series middle series hole 918d hole seg58 hole 917a imilik iii søndre aputitêq snout series lilloise kap boswell kræmer ø/ kangerlussuaq syenite gabbro dyke sill central intrusive complex lava succession alkaline mafic and felsic mixed tholeiitic and alkaline tholeiitic plateau basalts lower basaltspa le oc en e o d p le gs 1 52 +1 63 o d p le g 16 3x o d p le g 16 3 o d p le g 16 3x ka p g us ta v h ol m re gio n ki ali ne q re gio n n ua lik re gio n ka ng er lu ss ua q sc or es by s un d jam es on l an d tr ail l ø g eo gr . s oc ie ty ø bo nt ek oe ø h ol d w ith h op e sh an no n w ol las to n fo rla nd bl os se vil le ky st eo ce ne o lig oc en e m io ce ne mols bjerge sk jo ld un ge n borgtinderne fig. 5 radiometric ages of cenozoic igneous rocks along the east greenland margin (larsen et al. 2014). odp: ocean drilling program. transition from rift to drift is indicated at 55 ma. bj: bjerg. cgl: conglomerate. pr: prinsen. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 12 of 154 www.geusbul let in.org c b tb dr lr lbjlb sdb d a e structural high mesozoic basin oceanic crust 100 km 25 km 25 km 25 km 2 0 0 nw se imu imu w e w e 2 0 1 2 id bf 3 d f b c e a imu itbf w e tw t (s ) tw t (s ) tw t (s ) tw t (s ) tw t (s ) koldewey platform 0 1 2 3 4 5 6 0 1 2 3 4 5 6 danmarkshavn basin danmarkshavn ridge danmarkshavn ridge thetis basin danmarkshavn basin thetis basin 10 12 8 6 2 4 0 d ep th (k m ) 25 km nw jameson land basin liverpool land ridge liverpool land basin sea level se fault intrusions basement volcanic rocks devonian – middle permian upper permian – jurassic palaeogene cretaceous neogene aggrading sequence prograding sequence basement presumed paleozoic sediments presumed mesozoic sediments erosional incision faultearly eocene unconformity base palaeogene intra-miocene unconformity (imu) top upper prograding unit clinoform 25 km fig. 6 interpreted seismic profiles from the margin of north-east greenland. a: profile across the north-east greenland shelf (redrawn after petersen 2019). prograding wedges are visible above the early eocene unconformity and above the intra-miocene figure 6 continued on next page http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 13 of 154 www.geusbul let in.org greenland at 7 ma (larsen et al. 1994b). however, biomolecules extracted from silt at the base of deep ice cores in central southern greenland indicate that the region was forested within the past million years (willerslev et al. 2007). further, schaefer et al. (2016) showed that greenland was deglaciated for extended periods during the pleistocene epoch (from 2.6 ma to 11 700 years ago), based on measurements of cosmic-ray-produced isotopes in a bedrock core from central greenland. recently, christ et al. (2021) concluded that the greenland ice sheet persisted through much of the pleistocene but melted and reformed at least once since 1.1 ma based on investigations of sub-glacial sediments at the base of a thick ice core. the greenland ice sheet, which today has a thickness of up to 3 km, thus appears to have had a discontinuous history. 2.3 palaeogene intrusives a large number of intrusions and magmatic complexes were emplaced in east greenland during and after breakup in the north-east atlantic at the paleocene– eocene transition (see fig. 5 for an overview of igneous activity along the coast). several major intrusive complexes are present within the study area; from south to north they include (nielsen 2002; larsen et al. 2014): • werner bjerge complex and malmbjerget in the northern part of jameson land of oligocene age, c. 30 ma and 25 ma, respectively. • kap simpson and kap parry in the eastern part of traill ø of middle eocene age, c. 39 ma. • kap broer ruys on the south-eastern edge of hold with hope of middle eocene age, c. 47 ma. • myggbukta on the south-western edge of hold with hope with an inferred middle eocene age (c. 47 ma) based on the nearby intrusive complex at kap broer ruys. a late sheet is dated to be of early oligocene age, c. 33 ma. many areas are dominated by dykes and sills, including: • jameson land. dykes of early eocene age (c. 52 ma). • mols bjerge on traill ø. sills of early eocene age (c. 54 ma) and dykes of late eocene age (c. 36 ma). • hvalros ø, north-east of wollaston forland. dyke dated to early miocene (c. 20 ma). 2.4 sediments and unconformities offshore north-east greenland the north-east greenland margin is narrow off liverpool land (about 80 km wide) but broadens considerably to the north where it is over 300 km wide east of store koldewey (fig. 6). there is a dense array of seismic reflection data over the margin but only limited seismic refraction data. in addition, there is very little sampling or drilling of the offshore sediments. most of what is known about the pre-basalt strata in the offshore area is thus based on seismic interpretation with reference to onshore geology and comparison to the conjugate norwegian margins. hamann et al. (2005) and tsikalas et al. (2005) provided overviews of the margin based on seismic data collected in the early 1990s, and fyhn et al. (2012) provided an update based on all data available at that time. the narrowest part of the margin is heavily affected by palaeogene volcanism. deeper basins and structures are thus obscured and cannot be defined. to the north as the margin widens, the influence of volcanism decreases and the less volcanic region is broadly divided into the danmarkshavn and thetis basins, separated by a structural high, the danmarkshavn ridge. both basins can be subdivided into smaller basins and the ridge is a complex structure that becomes less pronounced to the north (hopper et al. 2014). the more landward danmarkshavn basin is located between the koldewey platform and the danmarkshavn ridge. both the danmarkshavn and thetis basins are marked by sill intrusions, but evidence for significant extrusive volcanism is only found seaward near the continent–ocean transition zone, south of the danmarkshavn and thetis basins (south of 76°n, hamann et al. 2005). the margin has experienced a number of rift episodes with intervening periods of thermal subsidence from the late palaeozoic up to final continental break-up and seafloor spreading by the eocene. the northern danmarkshavn basin, and parts of the thetis basin, include thick salt units that were mobilised and became diapiric in many parts of the basin. based on the age of evaporites onshore and comparison to the south-west barents sea, figure 6 continued unconformity (imu; middle to late miocene, 15–10 ma; døssing et al. 2016). the cenozoic succession is tilted across the danmarkshavn basin and danmarkshavn ridge and offset by faults. b: profile across the north-east greenland shelf (hamann et al. 2005; hopper et al. 2014). c: profile across jameson land and the liverpool land basin (larsen 1990; stoker et al. 2017). d, e: profiles across the north-east greenland shelf illustrating the imu within the cenozoic succession. dr: danmarkshavn ridge. idbf: intra-danmarkshavn basin fault. itbf: intra-thetis basin fault. jlb: jameson land basin. lb: liverpool land basin. lr: liverpool land ridge. sdb: southern danmarkshavn basin. tb: thetis basin. twt: two-way time. the locations of panels a–e are shown in panel f. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 14 of 154 www.geusbul let in.org the salt is presumed to be late carboniferous – permian in age. significant reflectivity is observed below the base salt, indicating that the devonian–carboniferous basins in the onshore areas are well expressed offshore as well. above the interpreted salt horizons, thick successions of mesozoic and cenozoic sediments are observed. the profiles in figs 6a, b do not however, cross the salt structures. deposits of inferred latest jurassic – early cretaceous age are present over wide areas, and the basal boundary of this unit is marked by one of the highest amplitudes and regionally most extensive horizons on the north-east greenland shelf (hamann et al. 2005). it shows a characteristic low-angle, erosional truncation of the underlying high-amplitude reflectors. hamann et al. (2005) correlated this unconformity to the mid-volgian unconformity (c. 145 ma) at the base of the wollaston forland group onshore and interpreted it to mark the onset of rifting in the latest jurassic. deposition in the danmarkshavn basin, off store koldewey, was controlled by a few major faults similar to those at the margins of the jameson land basin (fig. 6c), in contrast to the large number of smaller scale faults described in the northern part of onshore east greenland. thermal subsidence led to continued deposition of thick cretaceous successions prior to a second major rift episode that culminated with the final break-up and the onset of sea-floor spreading at the paleocene–eocene transition. although cenozoic deposition continued postbreakup, much of this has been removed, most recently by glacial erosion of the shelf. the nearshore parts of the margin in many places have only a thin quaternary cover with mesozoic sediments sub-cropping near the sea floor. two important phases of progradation away from north-east greenland occurred after break-up. the youngest of these prograding wedges are up 200 km wide and are bounded below by a distinct angular unconformity, the intra-miocene unconformity (imu; figs 6d, e; hamann et al. 2005; berger & jokat 2008; døssing et al. 2016; petersen 2019). berger & jokat (2008) dated the imu at c. 14 ma based on correlation between seismic lines and results from ocean drilling program (odp) site 913 and suggested that the onset of progradation may have been caused by rapid changes in sea level and/or glacial erosion by an early ice sheet or coastal glaciers. however, døssing et al. (2016) suggested that the age of the imu was middle to late miocene (15–10 ma). defining an age for the imu is difficult due to poor recovery of material around the imu at odp site 913 (myhre et al. 1995). the definition was later improved by including reanalyses of data from all seven sites of this leg of the odp (hull et al. 1996). the imu is associated with low-amplitude folds and reverse, margin-parallel faults, with throws up to 200 m, suggested to be related to mild compression (hamann et al. 2005). margin-parallel, open folds that occur on traill ø have likewise been interpreted as evidence of post-breakup compression, possibly in late cenozoic time (price et al. 1997). others have observed that the imu defines a 20 km wide, domal structure with associated faulting within the continent–ocean transition zone, and that the dome resulted from a mild compressional phase (døssing et al. 2016). the imu marks the termination of synrift deposition in the deep-sea basins and the formation of a continuous plate boundary along the mohns and knipovich ridges in the north-east atlantic, which leads to an accelerated widening of the fram strait (døssing et al. 2016). the imu event is linked to mild compression and to the onset of late miocene uplift and massive shelf progradation on the north-east greenland margin. the correlation between margin uplift and plate motion changes demonstrate that the uplift was triggered by plate tectonic forces (døssing et al. 2016). older prograding clinoforms are steep and occur above a horizon defining erosional incision (petersen 2019). this erosional horizon truncates an early eocene unconformity, which can be relatively well dated due to its association with volcanism at break-up, but the erosional horizon does not truncate the imu and cannot be dated directly. as a significant sedimentary succession is present between the erosional horizon and the imu, petersen (2019) dated the erosional horizon as late eocene – mid-miocene. the steeply prograding clinoforms are caused by extensive mass wasting into the outer part of the shelf (the thetis basin) due to uplift of the inner part of the shelf (the danmarkshavn basin and danmarkshavn ridge). the most likely source areas for this progradational event are the south danmarkshavn ridge, the danmarkshavn basin and the area onshore north-east greenland south of store koldewey (petersen 2019). petersen also found that the steepness of the incision, in combination with the chaotic nature of the material transported into the thetis basin resulting from the mass-wasting event, indicated a rapid and significant tectonic movement. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 15 of 154 www.geusbul let in.org large-scale landscapes, such as mountains, plains and valley systems, are the visible expressions of the longterm effects of major erosional and tectonic events (brunsden 1993; ahnert 1998). the analysis of largescale landscapes thus affords insights into the tectonic development of a region over long time spans (lidmar-bergström 1996; bonow et al. 2006a, b, 2014; green et al. 2013; lidmar-bergström et al. 2013). evidence of uplift and subsidence may be obtained from stratigraphic landscape analysis, which is based on observation of cross-cutting relationships between peneplains and stratigraphic constraints from the sedimentary cover that preserved them (lidmar-bergström et al. 2013). here we adopt the definition that all surfaces graded to a base level are peneplains, irrespective of whether their detailed form is hilly or flat (green et al. 2013; lidmar-bergström et al. 2017). the formation of a peneplain involves kilometre-scale erosion of rock during valley incision and valley widening by running water and slope processes, whereas the final shape of the peneplain depends on the climatic conditions during its formation. in this way, stratigraphic landscape analysis allows construction of a relative chronology of denudation episodes leading to surface formation and to identification of the corresponding tectonic events. bonow et al. (2014) mapped regional erosion surfaces in southern east greenland (68–71°n) and bonow & japsen (2021, this volume) extended this mapping to north-east greenland, south of 78°n. in this chapter, we summarise the results of these studies. 3.1 preand post-basalt peneplains elevated plateaus with deeply incised valleys characterise the large-scale landscape of north-east greenland. bonow et al. (2014) identified three peneplains in southern east greenland a pre-basaltic, etch – or weathered – basement surface (es; etch surface) and the post-basalt upper and the lower planation surfaces (ups and lps, respectively). bonow & japsen (2021, this volume) continued the work of bonow et al. (2014) by presenting coherent maps and analyses of the large-scale landscape of north-east greenland (north of scoresby sund, 70–78°n), where these surfaces are important markers for the tectonic development of the margin. the es extends across crystalline basement along scoresby sund (larsen et al. 1989). the surface is characterised by distinct hills, a few hundred metres high, and limited by narrow valleys along old fracture systems. on milne land, the surface is preserved below middle jurassic sediments and at different locations around scoresby sund, below the more widespread palaeogene basalts (birkelund & perch-nielsen 1976; larsen et al. 1989). this surface occurs at various elevations, from sea level up to 1800 m a.s.l. the sub-jurassic es is tilted along the east coast of milne land and cut off by the near-horizontal ups. the ups defines the high plateaus, where it cuts across palaeogene basalts and older rocks (fig. 7). the ups north of scoresby sund is a major surface at c. 2 km a.s.l. away from the coast, where it extends across an area up to 200 km wide. it is dominated by basement terrains but also includes basement covered by palaeogene basalts (bonow & japsen 2021, this volume). in the northern interior, the ups also cuts across the palaeogene basalts near the greenland ice sheet at about 2.2 km a.s.l. (arnold escher land; 74°n, 28°w). in the northernmost part of the study area the ups is at slightly lower elevation, at around 1.5 km, but reaches elevations >3 km south of scoresby sund (bonow et al. 2014). the ups is often covered by ice due to its high elevation, and in large areas north of 74°n the ups disappears below the greenland ice sheet. where the ice sheet becomes coherent, the ups is only visible as flattish remnants at the ice margin. the extensive ice covers that reach the coast prevent a clear correlation between the well-developed ups in the south and the remnants in the north. landscapes across the sedimentary basins along the coast are characterised by many ridges. here, the ups is interpreted to correspond to minor flattish summits at c. 2 km a.s.l., for example on western traill ø. at stauning alper, summits reach an elevation of 2.8 km a.s.l., the highest elevation within the study area. here, the contours of the ups define a domal pattern that can be traced up to 2.3 km a.s.l. and indicates a localised, tectonic movement after the formation of the ups. in parts of the blosseville kyst region, south of scoresby sund, there is little angular discordance between the ups and underlying basalt flows, but the discordance is pronounced in the inner parts of scoresby sund (larsen et al. 1989; pedersen et al. 1997). study of zeolite zones in the basalts by larsen et al. (1989) led to the conclusion that regional erosion removed the uppermost basalt units after tilting of the lava succession, and that 700 to 900 m of lavas had been removed by erosion across gåseland and milne land (fig. 9). the lps is incised below the ups and is thus younger (fig. 8). the lps is readily identified north of scoresby sund, particularly around milne land, where it forms valley benches at approximately 1 km a.s.l. below the ups (figs 8, 9). towards the coast, the valley benches merge 3. stratigraphic landscape analysis http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 16 of 154 www.geusbul let in.org b a fønfjord gåseland milne land gåsefjord geikie plateau 10 km basement milne land fm geikie plateau fm rømer fjord fm skrænterne fm cover vvvv v v v v v v nwse top of lava pile predicted from elevation of zeolite zones post-basalt offset ice surface c basement basement basalt 1.8 km es es ups ups ups 0 1 2 km fig. 7 the upper planation surface (ups) is well developed across the palaeogene basalts on gåseland and across basement rocks on milne land (background of the photo). a: view towards north-west across the inner part of gåsefjord. b: the same view with annotations. the basalts cover an undulating, deeply weathered basement surface; the etch surface (es). the elevation of ups towards the coast is c. 1.8 km a.s.l. c: profile showing the inclined and truncated strata of the main basalts and the original top of the lava pile estimated from the elevation of zeolite zones across milne land, gåseland and geikie plateau (larsen et al. 1989). the present-day lava surface is consequently an erosional feature and the tilt and offset of the basalt formations are post-basalt features. the ups on gåseland marked in b is thus about 700 m below the original top of the volcanic pile. note the tilt of four formations of the main basalts: geikie plateau fm, skrænterne fm, milne land fm and rømer fjord fm. the profile in c crosses gåseland a few kilometres west of where the photo was taken. location of photo and profile in fig. 9. reproduced from bonow & japsen (2021, this volume); based on larsen et al. (1989) and bonow et al. (2014). http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 17 of 154 www.geusbul let in.org to form a more coherent surface that extends across an up to 100 km wide zone. in the areas where sedimentary rocks crop out, the lps is defined by flat-topped summits, while in the areas with basement rocks, the lps is more extensive and less dissected. north of 74°n, the lps dominates the ice-free areas of the upper-plateau landscape in near-coastal areas. the elevation of the lps is relatively low in the northernmost parts of the study area, for example 600 m a.s.l. on germania land, but it reaches 2 km a.s.l. south of scoresby sund (bonow et al. 2014). the ups is well developed in basalts south of scoresby sund with no evidence of faulting. in contrast, north of scoresby sund, where the pdmf is a prominent feature, the ups is not preserved east of the fault. however, the lps a lps lps lps nordvest�ord ups lps ups ups lps ups stauning alper stauning alper flyver�ord flyver�ord ups b elevation (m) 10005000 1500 2000 ups nordvest�ord lps fig. 8 relationship between the upper and lower planation surfaces along nordvestfjord (ups and lps, respectively). a: lps along flyverfjord and nordvestfjord; view towards north-east across th. sørensen land towards nathorst land and stauning alper. b: 3d elevation model in approximately the same position as photo in a. the lps is developed as valley benches along the main fjords. note that some minor valleys stop to incise at the lps (arrows). map location in fig. 9. reproduced from bonow & japsen (2021, this volume). http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 18 of 154 www.geusbul let in.org lps is defined on both sides of the fault at about the same elevation. west of the fault, wide valley benches define the lps, and the surface continues at the same elevation east of the fault, where flat-topped summits define the surface. the step (escarpment) between the ups and the lps follow the incision along the major valleys, whereas the pronounced topographical step across pdmf is caused by the differences in lithology. the ups and lps dominate the appearance of the landscape and can be identified along the entire margin, despite intense fluvial and glacial erosion after the formation of the lps (fig. 9). 3.2 relative denudation chronology bonow & japsen (2021, this volume) concluded that the ups is a peneplain that represents a single erosional phase that affected the entire margin of east greenland between 68 and 78°n. the ups formed by erosion into the palaeogene basalts and the underlying crystalline basement, implying that the peneplain was graded to base level after the extrusion of the palaeogene basalts. the base level at that time was the level of the adjacent atlantic ocean. consequently, the present elevation of the ups is the result of uplift that affected the margin after the formation of the ups. south of scoresby sund fig. 9 extent and elevation of the upper and lower planation surfaces (ups and lps, respectively). the ups is typically 2 km a.s.l. south of 74°n at some distance from the coast. north of 74°n, the ups is only identified as small remnants close to the ice sheet. the lps, which typically is at 1 km a.s.l., formed by incision along valleys below the ups, but the lps along the valleys coalesces to form a coherent surface towards the coast. note the domal structure of the ups centred at stauning alper. this pattern suggests a tectonic doming after formation of the ups. the profile in fig. 7c extends about 25 km south of the map. modified from bonow & japsen (2021, this volume). ups dominates lps dominates clavering ø arnold escher land liverpool land gåseland flyverfjord nordvestfjord nathorst land renland th sørensen land store koldewey germania land ups and lps overlap lps elevation (km) ups elevation (km) scoresby sund 74°n 50 km 72°n 70°n 30°w 25°w 20°w 1.5 1.0 1.1 1.1 1.1 1.0 fig. 8 fig. 7c 1.2 1.1 1.1 1.3? 1.0 1.4 1.2 1.5 1.8 1.8 1.7 2.3 2.0 stauning alper fig. 7a, b http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 19 of 154 www.geusbul let in.org (c. 69°n), the volcanics of the middle miocene vindtop formation accumulated on the ups, which had therefore formed by middle miocene times (storey et al. 2004; bonow et al. 2014). a significant change in base level (uplift) terminated the development of the ups, leading to incision of valleys below the ups and to development of a new surface grading towards the new base level defined by the adjacent ocean. this led to the formation of the lps. as the difference in elevation between the ups and lps typically is about 1 km, the magnitude of rock uplift in this episode was 1 km. a final phase of uplift interrupted the development of the lps, leading to incision of valleys below the lps and to the formation of the present-day relief. the elevation of the lps is about 1 km a.s.l., and this implies that magnitude of rock uplift since the formation of the lps was 1 km. the timing for this event is uncertain based on landscape evidence alone, but the valley benches associated with the lps are fluvial in origin. this suggests that the lps existed prior to the onset of the major late cenozoic glaciations around 2.5 ma (thiede et al. 1998). glacial action contributed to the destruction of the ups and lps. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 20 of 154 www.geusbul let in.org afta is a method for determining thermal histories of rocks at temperatures generally less than 130°c. the method depends on analysis of radiation damage features, fission tracks, in either accessory apatite crystals separated from igneous rocks or detrital apatite grains from sedimentary rocks (green et al. 1989, 2013; green & duddy 2012). in this chapter, we describe the principles of afta and present new data from samples from outcrops in north-east greenland. these data are supplemented by new vr data from mesozoic sediments. we derive thermal-history solutions in terms of discrete episodes of cooling for each of the samples and produce a regional synthesis of ten regional episodes of cooling ranging from the carboniferous to the pliocene (appendix 1). we discuss the mechanisms of cooling and heating inferred from the afta data, based on the information derived from the variation in palaeotemperatures with elevation in vertical transects (vts; appendix 2). the origin of each of these episodes is discussed in detail in chapters 5 and 6. 4.1 afta and vr principles fission tracks in apatite crystals are generated continuously through time by spontaneous fission of uranium atoms, present as an impurity within the apatite lattice and typically at levels of parts per million. these tracks are revealed in a polished surface by etching with dilute acid. the number of tracks that intersect a unit area of an apatite grain surface (track density) depends on the uranium content of the grain and the time over which tracks accumulated, following the laws of radioactive decay and on the track length (see galbraith (2005) for a detailed exposition of the underlying method). if all tracks have the same length, unaffected by heating, then by counting the number of tracks in unit surface area and measuring the uranium content, the time over which tracks accumulated can be determined – a time referred to as a fission-track age, which is not a geological age. in practice, tracks in natural apatites have a distribution of track lengths, determined by the thermal history. therefore, a fission-track age does not indicate the timing of a specific event but must be interpreted in conjunction with the measured distribution of track lengths to obtain constraints on the underlying thermal history (green & duddy 2012). tracks are formed with an initial length within a narrow range, but once formed they shorten by a process known as annealing, at a rate dependant on the prevailing temperature. because temperature dominates over time in the kinetics of the shortening process, as the sample temperature increases, all tracks are reduced to around the same length regardless of when they were formed. this process is irreversible, so if the temperature drops, all tracks formed up to that time are effectively ‘frozen’ at the length attained at the maximum temperature. tracks which form after cooling are longer, due to the lower prevailing temperatures. thus, at the end of a history involving simple heating and cooling, a sample will contain two populations of tracks; shorter tracks formed up to the onset of cooling and longer tracks formed after cooling. the proportion of short to long tracks will reflect the time of cooling relative to the total time over which tracks have been retained, while the mean length of the shorter component of tracks reflects the maximum palaeotemperature attained prior to cooling (green & duddy 2012). because the probability of tracks intersecting the surface (and contributing to the fission-track age) depends on track length, the shorter population of tracks will contribute a smaller proportion to the fission-track age compared to the longer tracks. therefore, the fission-track age underestimates the time over which tracks have accumulated. at some critical palaeotemperature the damage constituting the track is totally repaired, and the length is reduced to zero, referred to as total annealing. the precise temperature at which this occurs depends on the heating rate and the chlorine content of the apatite, but is typically in the range 100 to 130°c. in samples that reached higher palaeotemperatures prior to cooling, tracks are only retained after cooling below this limit. afta data from such samples provide only a minimum estimate of the maximum palaeotemperature, but typically provide good constraints on the timing of the cooling episode. throughout this manuscript, we use the terms ‘palaeothermal episode’ and ‘cooling episode’. palaeothermal episode refers to the time when the samples were hotter than they are now, whereas cooling episode refers specifically to the subsequent time when cooling began. figure 10 illustrates the development of fission tracks in an apatite crystal through a notional history involving initial cooling from >110°c at time = t1 to 10°c at t2, representing exhumation of basement to the surface. tracks formed during this initial cooling phase will have a range of lengths, as they experienced different maximum temperatures during cooling (situation e2 in fig. 10). tracks formed during reheating (representing reburial) between t2 and t3 are progressively shortened until the onset of cooling at t3. because the temperature does not subsequently exceed the maximum value reached at t3, all tracks formed up to this point retain 4. apatite fission-track analysis data http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 21 of 154 www.geusbul let in.org the length they attained prior to cooling. the sample is reheated (reburied) towards the peak temperature at t4, shortening tracks formed after t3. when cooling begins again at t4, these tracks remain longer than those formed prior to the earlier cooling event at t3. thus, the shorter tracks formed prior to cooling at t3 preserve evidence of the earlier cooling episode. after cooling at t4, temperature remains low and tracks formed between t4 and t5 are longer than all tracks formed previously. the track-length distribution measured at the present wt% chlorine fi ss io ntr ac k ag e (m a) 300 240 180 120 60 0 time (myr bp) time (myr bp) 1.0 0.8 0.6 0.4 0.2 0.0 0200300 300 240 180 120 60 0 1 2 3 4 5 le ng th re du ct io n l/ lo time (myr bp) 0 20 40 60 80 100 120 0200300 100 1 2 3 4 5 te m pe ra tu re (° c ) 1 2 4 5 3 40 30 20 10 0 0 5 10 15 20 track length (μm) n um be r o f t ra ck s a b c d e 100 t₁: 300 ma, t = 110°c. no tracks present. t₂: 275 ma, t = 10°c. tracks formed during cooling have di�erent lengths, re�ecting di�erent maximum temperatures. t₃: 125 ma, t = 90°c. more tracks formed, and all those formed since t₂ (blue) are shortened to a length determined by the maximum temperature at t₃. most of the tracks formed prior to t₂ are also shortened to the same length; only tracks that were shorter than this length remains with their length at t₂. t₄: 25 ma, t = 60°c. a new population of tracks formed between t₃ and t₄; all these are shortened to a length determined by the temperature at t₄, and are all longer than the tracks formed between t₂ and t₃. t₅: 0 ma, t = 10°c. a few tracks formed between t₄ and t₅ are longer than all other tracks in the sample. fig. 10 schematic illustration of the afta method (green & duddy 2012; green et al. 2013). a: schematic thermal history illustrating three periods of exhumation and two intervening periods of burial; points 1 to 5 corresponding to t1 to t5 (time = tn). b: shortening trajectories for tracks formed at different times through the history. tracks formed during different segments of the history are coloured-coded corresponding to the images in panel e (1 to 5). c: evolution of apatite fission-track age with time through the thermal history shown in a. the fission-track age initially evolves linearly with time (dashed line). as t3 approaches, the temperature nears 90°c causing length reduction (see b) resulting in significant age reduction. after cooling, length reduction is less severe and age evolution approaches linearity again. at t5 the measured age is less than the total duration of the history, due largely to length reduction at t3. d: resulting distribution of track lengths, reflecting the presence of three populations of tracks, shortened to differing degrees as shown in b. the natural spread in track length is c. ±1 μm within a single population of tracks. tracks belonging to the separate populations shown in b cannot be resolved, so individual tracks cannot be attributed to unique populations. e: schematics of tracks in a notional apatite grain through the thermal history in a. tracks colour-coded as in b. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 22 of 154 www.geusbul let in.org day (fig. 10d) is dominated by two populations of tracks: shorter tracks formed up to t3 and longer tracks formed between t3 and t4, plus a small proportion of very long tracks formed after t4. the proportion of these longest tracks is very small because cooling was recent, and these cannot be easily resolved in the distribution (fig. 10d). the final fission-track age (t5) is reduced due to shortening of the tracks formed before t3. peak palaeotemperatures and the time at which cooling began in the two episodes (t3 and t4) can be determined using the relative proportions of longer and shorter tracks in the track-length distribution, in conjunction with the fission-track age (see appendix 1 for further explanation). in addition, the history of the apatites in the sediment provenance terrain (t1 to t2) may be determined, if the temperature at t3 was not sufficiently high to overprint this part of the thermal history. where heating is predominantly due to depth of burial, palaeotemperatures determined from afta in isolated outcrop samples can be converted to former depth of burial, and hence amounts of removed section, using an assumed palaeogeothermal gradient. where samples are analysed over a range of depths in a borehole or over a range of elevations in mountainous terrain, the allowed range of palaeogeothermal gradients can be determined directly. in principle, this allows more precise determination of amounts of removed section. however, this analysis is subject to a range of assumptions (green & duddy 2012; green et al. 2013) and results should be viewed with caution. 76 °n 74 °n 72 °n 70 °n 70 °n 72 °n 74 °n 76 °n 30°w 20°w 30°w 20°w 19 8 7 23 20 48 3 38 41 40 39 16 17 1815 8 9 14 13 9 3 10 4 5 10 11 1 11 2 6 73 72 202-12 202-11 103-3 27,28 103-4 159-1 103-6 159-2 202-1 202-9 1019-32 21,22 103-1,2 159-4 159-3 7 4 47 46 41 5 6 174 9 3 1 10 11 13 12 8 15,16 65 24-26,70 42-45 202-3,4,5,6,7,16 49-55 29-33 35-37 63-64 57-62 22-26 27-32 89-92 33,34,35,36 37,38,39 18,20,21 42 47 49 5051 52,54,55 5657 58 59 85 84 83 79 73 74 72 75 71 70 60 61 7776 78 62 66 6567 68 69 63 64 88 101,102 103 104 105 100 98,99 86,87 80,81,82 43-46 93,94,95, 96,97 66-69 1 23 4 4a 5 6 7 8 9 15 3 62 vertical transect gc1016 sample 15 gc1104 sample gc1077 sample gc522 sample 1 9 10 14 11 12 13 15 17 16 50 km 4 5 6 7,8 9 21 10 12 13 14 15 19 16,17 18 24: nanok-1 b/h 1 2 3 20 22,23 pre-devonian rocks fault devonian sediments carboniferous–cretaceous sediments palaeogene basalts palaeogene intrusives cenozoic sediments quaternary sediments ice sea fig. 11b fig. 11c a 2 3 4 6 5 1 inland north inland south central inland liverpool land jameson land traill ø geographical society ø germania land milne land ymer ø outlier fig. 11 sample locations. a: afta samples across the entire study area. one sample from a parallel study (gc101932; japsen et al. 2014), for which data are not presented here, is also shown. red lines indicate the six regions (numbered 1–6) referred to in table 1. b: afta and vr samples on store koldewey. c: afta and vr samples in the clavering ø area. df: dombjerg fault. pdmf: post-devonian main fault. geology based on fig. 3. arrow at the northern edge of the map points to the location of sample gc107711 just north of the map frame. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 23 of 154 www.geusbul let in.org vr data provide a useful complement to afta data in providing independent estimates of maximum post-depositional palaeotemperatures. vr is based on the reflectance of coaly material in sediments, which increases at a rate that depends on temperature. the kinetics of this process are well documented (sweeney & burnham 1990) and are similar to those of fission-track annealing in apatite (see green & duddy 2012). integration of afta and vr allows data to be collected from different lithologies within the same succession and provides an independent check on maximum post-depositional palaeotemperatures obtained from either method. 8 7 23 20 48 3 38 41 40 39 24-26,70 49-55 24-26, 70 27-32 63-64 57-62 22-26 27-32 89-92 47 49 85 84 79 73 74 72 75 70 88 105 100 98,99 86,87 80,81,82 43-46 93,94,95, 96,97 66-69 1 2 3 4 4a 9 10 14 16 24: nanok-1 b/h 25 km kuhn ø clavering ø vr sample (gc1077) dombjerg d f pd m f hudson land payer land giesecke bjerge wollaston forland hope with hope 74 °n 74 °n 17 13 12 83 101,102 22,23 store koldewey c 21°w 21°w b 19°w 19°w 15,16 103 104 20 15 km 76 °1 5' n 76 °1 5' n 109.1 139.1 144.1 145.1 110.1 111.1 112.1 116.1 107.1,108.1 113.1, 114.1, 115.1 125.1 126.1 106.1 140.1 119.1, 120.1, 121.1 117.1 118.1 123.1 122.1,124.1 141.1 127.1, 128.1, 129.1 142.1, 143.1 133.1, 134.1, 135.1 130.1, 131.1, 132.1 136.1, 1371, 138.1 middle–upper jurassic vr sample (gc1077) lower cretaceous fig. 11 (continued) sample locations. a: afta samples across the entire study area. one sample from a parallel study (gc1019-32; japsen et al. 2014), for which data are not presented here, is also shown. red lines indicate the six regions (numbered 1–6) referred to in table 1. b: afta and vr samples on store koldewey. c: afta and vr samples in the clavering ø area. df: dombjerg fault. pdmf: post-devonian main fault. geology based on fig. 3. arrow at the northern edge of the map points to the location of sample gc1077-11 just north of the map frame. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 24 of 154 www.geusbul let in.org 4.2 previous apatite fission-track studies previous apatite fission-track studies in north-east greenland provide important insights into the thermal and tectonic history of the region. early studies focussed on the palaeogene intrusions along the margin south of scoresby sund (e.g. gleadow & brooks 1979), detailing the post-intrusion cooling of these bodies. further studies of basement terrains and sedimentary successions of jameson land revealed major cenozoic cooling along the margin, interpreted in terms of kilometre-scale denudation related in some way to the process of rifting and separation of the north atlantic ocean (hansen 1988, 1992, 1996). thomson et al. (1999) published afta results from the sedimentary successions to the north of jameson land (sample locations shown in fig. 11 with the prefix gc522), reporting two cenozoic cooling episodes which began between 40 and 30 ma (late eocene – oligocene) and between 10 and 5 ma (late miocene). an earlier episode of cooling, which began sometime between 225 and 165 ma, was also recognised in inland samples. the late eocene – oligocene episode was interpreted to be due either to uplift and erosion or to hydrothermal effects associated with igneous bodies of that age. the late miocene episode, in which several samples cooled from palaeotemperatures as high as 90 to 100°c, was interpreted as due to uplift and erosion related to a change in the north atlantic spreading vector. thomson et al. (1999, p. 1045) noted that “no palaeo-thermal effects have been identified related to the onset of rifting in the early tertiary”. johnson & gallagher (2000) interpreted results from samples of carboniferous sandstones from a vertical transect on clavering ø as defining progressive cooling through late palaeozoic and mesozoic times in a series of stages prior to accelerated cenozoic cooling. the section sampled by johnson & gallagher (2000) is overlain by palaeogene basalts, showing that their uppermost sample must have been at near-surface temperature in the palaeogene. however, their analysis took no note of this basic constraint on the thermal history of the succession, leading to serious underestimation of final mesozoic cooling and palaeogene re-burial prior to the onset of cenozoic cooling (japsen et al. 2010; green et al. 2013). johnson & gallagher (2000) reported no evidence of palaeogene palaeothermal effects. hansen et al. (2001) reported results from the jameson land basin and adjacent regions, showing a consistent regional pattern with youngest ages less than 25 ma restricted to a region around the north and south coasts of kong oscar fjord, surrounded by a wider region characterised by ages between 25 and 45 ma. older ages were largely restricted to higher elevations and locations to the south and west of jameson land. hansen et al. (2001) interpreted these results as reflecting a complex interplay between deeper burial prior to cenozoic exhumation, circulation of hot fluids associated with mid-cenozoic intrusive activity around the eastern end of kong oscar fjord and possibly elevated basal heat flow in eastern traill ø. hansen et al. (2001) favoured an onset of exhumation at c. 55 ma in the south of the basin, and before 25 ma in the north, suggesting removal of as much as 4–5 km for palaeogeothermal gradients around 25°c/km. the apatite fission-track ages presented by hansen et al. (2001) show an erratic variation, suggesting significant local heating effects. in view of these observations, any evidence for palaeogene exhumation in their dataset is tentative at best. palaeogene palaeothermal effects thus appear to be restricted either to contact heating from intrusions and associated hydrothermal effects. apatite (u-th)/he ages of basement rocks reported by swift et al. (2008) from locations close to those studied previously by thomson et al. (1999) and johnson & gallagher (2000), showed erratic variation with elevation. these results were interpreted as indicating a phase of rapid denudation, beginning 74 ± 15 ma. this is a stark contrast to previous studies in the region, which showed no evidence of regional late cretaceous – palaeogene cooling. swift et al. (2008) interpreted their (u-th)/he ages in terms of cooling below 35–75°c. however, subsequently, abundant evidence has emerged that the thermal response of the apatite (u-th)/he system is more complicated, and that a range of external factors affect the measured ages (green & duddy 2018). for this reason, no significance can be attached to the conclusions of swift et al. (2008). this is further emphasised by results from further south on the east coast of greenland presented by hansen & reiners (2006), who reported that thermal histories indicated by apatite fission-track and (u-th)/he data were inconsistent, and the he-ages should be discounted. japsen et al. (2014) interpreted afta data in 90 samples from southern east and south-east greenland, primarily between 68 and 70°n, in terms of nine discrete cooling episodes since the late palaeozoic. some of these episodes were regional while others were more localised. japsen et al. (2014) combined the interpretation of the afta data with results from stratigraphic landscape analysis (bonow et al. 2014) and found that three regional phases of uplift and exhumation shaped the present-day margin: (1) a late eocene phase of uplift led to formation of a regional erosion surface, the ups, near sea level. (2) late miocene uplift led to formation of another erosion surface, the lps, by incision below the uplifted ups. (3) pliocene cooling could not be resolved in the afta data, but a pronounced pliocene phase of uplift and incision was established from landscape http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 25 of 154 www.geusbul let in.org evidence (bonow et al. 2014), which led to the final creation of the present-day topography of the region. thus, despite several regional thermal-history studies of the east greenland margin, opinions differ with regard to the precise time at which individual rock units reached their palaeothermal maximum, how this timing varies across the region, and the nature of the mechanisms responsible for heating and cooling. 4.3 new afta and vr data afta has been applied to 217 outcrop samples of basement and sedimentary rocks in north-east greenland between 70 and 78°n (fig. 11). methods are described in green et al. (2013). methods for sample preparation, analytical details, basic afta data in individual samples and the thermal-history interpretations derived from afta data in each sample are reported in three geotrack company reports and are available as supplementary material: • report gc1016 (95 samples; green 2009; supplementary file s1) • report gc1077 (99 samples; green 2011; supplementary file s2) • report gc1104 (23 samples; green 2012; supplementary file s3). all sample details (appendix 1.2), afta data (appendix 1.3) and thermal-history interpretations (appendix 1.4) are available in supplementary file s4. supplementary file s5 contains sample coordinates and thermal-history solutions in kmz format. results in 11 outcrop samples from traill ø, originally published by thomson et al. (1999), are included in geotrack report gc1016 (samples with prefix gc522). these samples were reanalysed and reported in geotrack report gc1016 to incorporate measurement of wt% cl, providing updated thermal-history interpretations for these samples. both the basic afta parameters and the interpretations are similar to those quoted by thomson et al. (1999). apatite separates of 20 palaeozoic and mesozoic sedimentary rocks from the jameson land region were kindly made available by kirsten hansen (samples with prefix gc103, gc159, gc202). these samples were first analysed in the 1980s by geotrack for atlantic richfield company (arco) and results were published by hansen et al. (2001). the samples were recounted and re-interpreted using modern techniques and included in geotrack report gc1016. coordinates of samples from previous studies were digitised from published maps (thomson et al. 1999; hansen et al. 2001). while we have used these elevations in appendix 2, sample elevations derived in this way were found to conflict with values quoted by hansen et al. (2001; particularly for samples from schuchert dal, western jameson land), so some doubt surrounds these data. rocks sampled in this study comprise precambrian and caledonian basement, palaeozoic intrusions, palaeozoic and mesozoic sedimentary rocks and a small number of cenozoic sedimentary and intrusive rocks. results from two samples of palaeogene sediments at kap dalton (69.5°n), and from one sample (gc1019-32) of palaeogene sediments from kap brewster were published by japsen et al. (2014). apatite yields were generally excellent. while some samples provided lower amounts of apatite than generally required for an analysis of the highest quality, interpretations derived from such data are still reliable within the stated uncertainties (95% confidence limits), although these will be broader than those from higher quality data. nevertheless, results from such samples are still useful when assessed alongside higher quality data. samples collected at different elevations within restricted areas constitute 17 near-vertical transects (vts; appendix 2.1), to quantify palaeogeothermal gradients and amounts of removed section. geotrack report gc1077 (green 2011; supplementary file s2) also contains results of vr analyses undertaken by keiraville konsultants, australia, of 40 samples of outcropping sedimentary units from locations in the clavering ø area and on store koldewey (figs 11b, c). vr data, including maximum palaeotemperatures and the associated thermal episodes to which they are attributed are listed in appendix 1.5 (and in supplementary file s4). the variation in measured apatite fission-track ages across the region is shown in fig. 12. such a map can be misleading because measured fission-track ages are not only sensitive to the thermal history but also to depositional ages as well as external factors such as the composition of apatites within each sample and the sample elevation. regional variation is better addressed by investigating the palaeotemperatures in individual episodes estimated from the thermal-history interpretation of the data, as discussed in section 4.4. however, some broad generalisations are possible based on fig. 12. youngest fission-track ages (<25 ma) occur in samples from eastern traill ø and north-eastern jameson land, on either side of kong oscar fjord. fission-track ages increase north, west and south of these areas. away from the focus of young ages, a background pattern is evident where locations close to sea level typically yield ages around 100 to 200 ma while samples at highest elevations yield ages >200 ma. elsewhere, fission-track age variation with elevation is more erratic, for example, on north-west jameson land where ages show little correlation with elevation. the variation in thermal history within the vertical sample transects is discussed in http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 26 of 154 www.geusbul let in.org greater detail in section 4.6 and in appendix 2. for now, we simply note a complex variation in thermal history across the area, and in some cases at a local scale. mean track lengths measured in each sample are plotted against fission-track age in fig. 13a alongside published data. none of the fission-track ages are older than c. 300 ma, showing that the data are dominated by the phanerozoic history. the data define a broadly consistent pattern, although this is rather more complex than the simple ‘boomerang trend’, which characterises a suite of samples affected by a single dominant palaeothermal episode (green 1986). the more complex relationships in fig. 13a can be understood in terms of the superposition of a series of separate ‘boomerang trends’ (fig. 13b), suggesting that the samples have undergone a series of palaeothermal episodes, with different samples being heated to different temperatures in each episode. 4.4 thermal-history solutions from afta thermal-history solutions derived from the afta data in each sample (using the approach described in appendix 1) are provided in appendix 1.4. solutions are presented as intervals representing 95% confidence limits on the 30°w 20°w 30°w 20°w 100 km 76 °n 74 °n 72 °n 70 °n 76 °n 74 °n 72 °n 70 °n fission-track age (ma) <25 25–45 45–60 60–100 100–200 >200 0.0 0.2 0.4 0.6 0.8 1.0 1.2 1.4 el ev at io n (k m ) fig. 12 fission-track ages in individual outcrop samples. rectangular symbols refer to vertical profiles. significant regional variation is evident, with youngest ages around the eastern end of traill ø and north-east jameson land. arrow at the northern edge of the map points to the location of sample gc1077-11 just north of the map frame. geology as in fig. 11. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 27 of 154 www.geusbul let in.org maximum palaeotemperature and timing of cooling (as defined in appendix 1.1) from one, two or three palaeothermal episodes (i.e. periods when samples were hotter than they are today). the number of palaeothermal episodes that can be defined in each sample depends on how well separated the events are in temperature and time, as well as the number of track lengths and fission-track ages measured in each sample. estimates of the timing of discrete cooling episodes in individual samples are compared in appendix 1.6 for six regions along the coast, illustrated in fig. 11a (these regions are defined simply to present the data in convenient sub-sections). based on the general uniformity of the data, we have sought to define the minimum number of regionally synchronous cooling episodes, which can explain the results in all samples. where palaeotemperatures and estimates of the onset of cooling are similar in adjacent samples, the constraints are interpreted, as representing a common episode. in this way, samples with quite broad constraints can still be allocated to specific regional events. as an example, all samples from vt1 are interpreted as defining the earliest cretaceous c3 cooling episode, even though the timing constraints for this episode are quite broad in some samples (for example, 320–60 ma, sample gc1016-50). however, the constraints on the c3 palaeotemperatures for vt1 are tightly constrained and show consistent variation with elevation (appendix 2.2), and it is thus reasonable to assume that all vt1 samples cooled in episode c3. this illustrates how data with quite broad constraints can provide useful conclusions. the timing of regionally synchronous cooling episodes in the six regions are presented in table 1. in many cases the timing of these major cooling episodes is consistent between different regions. in a few cases, cooling is seen in some regions and not others. thus, 9 10 11 12 13 14 15 16 0 50 100 150 200 250 300 350 m ea n tr ac k le ng th (m ic ro ns ) fission-track age (ma) 9 10 11 12 13 14 15 16 m ea n tr ac k le ng th (m ic ro ns ) a b 0 50 100 150 200 250 300 350 fission-track age (ma) thomson et al. (1999) hansen et al. (2001) gc1016 johnson & gallagher (2000) gc1077 gc1104 late palaeozoic to early mesozoic re-setting of early palaeozoic histories early cenozoic re-setting of palaeozoic histories late cenozoic re-setting of earlier histories fig. 13 relationship between mean track length and fission-track age. a: outcrop samples from north-east greenland. the new data span the range of previous studies. b: same as a, overlain with coloured lines to indicate the cumulative effects of resetting by several separate palaeothermal episodes, each of which has affected individual samples to varying degrees. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 28 of 154 www.geusbul let in.org we can distinguish between regional and local episodes. ten cooling episodes (c0–c9) are recognised within the study area. the onset of cooling in each episode ranges from the late palaeozoic to the early pliocene (table 1). 4.5 regional variation in the magnitude of palaeothermal episodes maps of palaeotemperatures characterising the ten palaeothermal episodes that preceded the cooling episodes (c0 to c9) are shown in fig. 14. palaeotemperatures derived from vr data from 40 samples presented in geotrack report gc1077 are also included where data are available (mainly from locations in the clavering ø area and on store koldewey). results from the vr data are attributed to regional events based on palaeotemperatures from afta in nearby samples (appendix 1.5). detailed palaeotemperature maps for store koldewey and for the region around clavering ø are shown in figs 15 and 16, respectively. the palaeotemperature maps illustrate how some events only affected restricted areas, while others affected much of the region. here we discuss general trends within these data. evidence for mechanisms of heating and cooling in each episode provided by the variation of palaeotemperatures with elevation in vertical transects is discussed in section 4.6. it is clear from these maps that many areas have undergone several regional cooling episodes. because only three events can typically be resolved from afta data in a single sample, this can cause difficulties in resolving discrete episodes in individual samples, particularly where events that are close in time are of similar magnitude. thus, in some cases, the quoted thermal-history solutions may represent the combined effects of two or more unresolved multiple cooling episodes. this is particularly important also at low temperatures (<70°c), where differences in track-length reduction caused by separate palaeothermal episodes may be insufficient to allow multiple episodes to be resolved. however, the generally consistent results discussed here suggest that this does not constitute a serious problem in interpretation. a few cases where this problem may be acute are discussed. 4.5.1 late carboniferous episode (c0) cooling began between 320 and 300 ma. episode c0 is recognised mainly in basement samples from the north and south of the region (c0 in fig. 14). maximum palaeotemperatures in this episode are in excess of 100°c for all samples in which this episode is recognised. since this is the earliest of the palaeothermal episodes identified across the region, it is only recorded in areas where palaeotemperatures in subsequent episodes have remained sufficiently low, such that evidence of this episode was not overprinted. locations that record episode c0 therefore represent relatively stable areas, which have undergone less post-permian heating and cooling than surrounding regions. note that the small number of samples that preserve evidence of the c0 episode emphasises the magnitude of post-palaeozoic palaeothermal effects in the study area, which have probably overprinted the effects of this episode over most of the region. regions defined in figure 11a. also shown are possible origins of the cooling during each episode. region 1: germania land, store koldewey, inland north. region 2: clavering ø, wollaston forland, kuhn ø. region 3: hold with hope, hudson land, ymer ø, geographical society ø. region 4: inland (central and south), milne land. region 5: traill ø, stauning alper, jameson land north, schuchert dal. region 6: jameson land south, liverpool land, blosseville kyst. appendix 1.6 displays the timing of discrete cooling episodes in individual samples. table 1 intervals defining the onset of cooling episodes (ma) based on afta data in 217 outcrop samples event region 1 region 2 region 3 region 4 region 5 region 6 synthesis onset of cooling (chronostratigraphy) mechanism c0 325–300 400–250 320–260 350–270 320–300 late carboniferous regional uplift and erosion c1 245–200 250–240 240–220 270–240 245–215 300–215 c. 240 middle triassic regional uplift and erosion c2 205–180 185–175 200–165 185–180 220–165 180–170 c. 180 early jurassic regional uplift and erosion c3 165–135 150–140 155–130 145–140 145–140 earliest cretaceous regional uplift and erosion c4 105–80 100–90 120–75 110–90 95–65 95–90 mid-cretaceous regional uplift and erosion c5 65–55 62–42 55–45 60–48 c. 55 early eocene local effects of intrusions c6 38–35 40–35 38–30 37–25 37–35 end-eocene regional uplift and erosion local effects of intrusions c7 20–18 25–18 early miocene local effects of intrusions local uplift and erosion c8 10–5 c. 10 10–7 15–10 12–10 12–10 c. 10 late miocene regional uplift and erosion c9 25–0 5–0 c. 5 c. 5 early pliocene incision of relief during regional uplift http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 29 of 154 www.geusbul let in.org fig. 15 c3 145–140 ma earliest cretaceous c2 c. 180 ma early jurassic c1 c. 240 ma middle triassic palaeotemperature (°c) >110 90–100 80–90 70–80 60–70 <60 100–110 palaeotemperature (°c) >110 90–100 80–90 70–80 60–70 <60 100–110 palaeotemperature (°c) >110 90–100 80–90 70–80 60–70 <60 100–110 palaeotemperature (°c) >110 90–100 80–90 70–80 60–70 <60 100–110 c0 320–300 ma late carboniferous 30°w 20°w 30°w 20°w 76 °n 74 °n 72 °n 76 °n 74 °n 72 °n 70 °n 30°w 20°w 30°w 20°w 76 °n 74 °n 72 °n 76 °n 74 °n 72 °n 70 °n 30°w 20°w 30°w 20°w 76 °n 74 °n 72 °n 76 °n 74 °n 72 °n 70 °n 30°w 20°w 30°w 20°w 76 °n 74 °n 72 °n 76 °n 74 °n 72 °n 70 °n fig. 15 fig. 15 fig. 15 100 km sa f pd m f d f milne land sa sa sa store koldewey kong oscar fjord clavering ø 100 km milne land store koldewey kong oscar fjord kuhn ø clavering ø wollaston forland 100 km milne land traill ø vt12 vt17 vt8 62 66 83 70 11 55 vt17 104 71 67 68 58 59 65 store koldewey kong oscar fjord kuhn ø clavering ø wollaston forland liverpool land liverpool land hudson land 100 km milne land fig. 16 store koldewey kong oscar fjord clavering ø germania land fig. 14 caption on page 31. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 30 of 154 www.geusbul let in.org c7 20–18 ma early miocene myggbuktaymer ø vt6 c6 37–35 ma end-eocene c5 55–45 ma early eocene c4 95–90 ma mid-cretaceous palaeotemperature (°c) >110 90–100 80–90 70–80 60–70 100–110 palaeotemperature (°c) >110 90–100 80–90 70–80 60–70 100–110 palaeotemperature (°c) >110 90–100 80–90 70–80 60–70 100–110 palaeotemperature (°c) >110 90–100 80–90 70–80 60–70 <60 <60 <60<60 100–110 30°w 20°w 30°w 20°w 76 °n 74 °n 72 °n 76 °n 74 °n 72 °n 70 °n 100 km milne land store koldewey kong oscar fjord kuhn ø clavering ø wollaston forland sa sa sø 30°w 20°w 30°w 20°w 76 °n 74 °n 72 °n 76 °n 74 °n 72 °n 70 °n 100 km jameson land traill økong oscar fjord clavering ø wollaston forland hold with hope liverpool land 30°w 20°w 30°w 20°w 76 °n 74 °n 72 °n 76 °n 74 °n 72 °n 70 °n 100 km jameson land liverpool land traill økong oscar fjord clavering ø 30°w 20°w 30°w 20°w 76 °n 74 °n 72 °n 76 °n 74 °n 72 °n 70 °n 100 km jameson land traill økong oscar fjord kuhn ø mestersvig mb sa sd 100 wegener halvø fig. 16 fig. 16 kap brewster hvalros ø germania land fig. 15 fig. 14 (continued) caption on page 31. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 31 of 154 www.geusbul let in.org 4.5.2 middle triassic episode (c1) cooling began at c. 240 ma. episode c1 is recognised sporadically in samples across much of the region (c1 in fig. 14). maximum palaeotemperatures in this episode are also in excess of 100°c for all samples in which this episode is recognised. around clavering ø, the c1 episode is recognised in many samples east of the pdmf, but one sample west of that fault records c1 palaeotemperatures in excess of 130°c because of high chlorine content in the apatites in this sample (gc107770). thus, it seems likely that c1 cooling affected areas both west and east of the fault, but samples west of the fault generally cooled below 110°c in cooling episodes after c1. as with the c0 episode, middle triassic cooling is only recorded in areas where palaeotemperatures in subsequent episodes have remained sufficiently low as to not overprint evidence of this episode. 4.5.3 early jurassic episode (c2) cooling began c.180 ma. episode c2 is recognised in many locations across the study area, in samples of basement and early palaeozoic sedimentary rocks, particularly where lower jurassic sediments are absent (e.g. milne land, traill ø, clavering ø and store koldewey; c2 in fig. 14). as such it is not recognised in coastal locations south of clavering ø, except for vt12 on liverpool land (appendix 2.3). maximum palaeotemperatures in this episode are around 100°c or above in most of the samples in which this episode is recognised, with a tendency towards lower values at the farthest c9 c. 5 ma early pliocene c8 c. 10 ma late miocene palaeotemperature (°c) >110 90–100 80–90 70–80 60–70 <60 100–110 palaeotemperature (°c) >110 90–100 80–90 70–80 60–70 <60 100–110 palaeotemperature from afta data palaeotemperature along vertical transect from afta data palaeotemperature from vr data gc1016 sample gc1077 sample 30°w 20°w 30°w 20°w 76 °n 76 °n 74 °n 72 °n 70 °n 30°w 20°w 30°w 20°w 76 °n 74 °n 72 °n 74 °n 72 °n 76 °n 74 °n 72 °n 70 °n 100 km jameson land jameson land milne land traill øtraill ø sd store koldewey kuhn ø clavering ø wollaston forland liverpool land 100 km 32 wegener halvø fig. 16 kong oscar fjord kong oscar fjord germania land 55 68 fig. 15 fig. 14 (continued) maps of palaeotemperatures derived from afta data. individual samples are attributed to palaeothermal episodes c0–c9 (table 1) and thermal history constraints are listed in appendix 1.4. detailed maps of palaeotemperatures derived from afta and vr data for store koldewey and the area around clavering ø are shown in fig. 15 and fig. 16, respectively. df: dombjerg fault. mb: malmbjerget. pdmf: post-devonian main fault. sa: stauning alper. saf: stauning alper fault. sd: schuchert dal. sø: sabine ø. vt: vertical transect. arrow at the northern edge of the map points to the location of sample gc1077-11 just north of the map frame. sample locations in fig. 11. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 32 of 154 www.geusbul let in.org inland locations and at higher elevations. this episode is only recorded in areas where palaeotemperatures in later episodes have remained sufficiently low, so as not to overprint evidence of this episode, which probably accounts for the absence of episode c2 in coastal and near coastal locations (c5 and c6 in fig. 14). 4.5.4 earliest cretaceous episode (c3) cooling began between 145 and 140 ma. episode c3 is recognised most prominently in the region around the western side of clavering ø, as well as in several samples across the interior of the study area (c3 in fig. 14). this contrasts with the early jurassic c2 episode, which is recognised at inland locations across the entire study area. palaeotemperatures characterising this episode are more variable than those identified in earlier events, although values are again in excess of 100°c at many locations, precluding definition of earlier events. palaeotemperatures also show a progressive decrease with increasing elevation in vertical transects which is particularly well displayed in data from clavering ø (vt1; appendix 2.3). offsets by faulting may occur between some of the inland sample locations. whereas some samples cooled from >110°c in this episode (e.g. gc107771, -67, -68), neighbouring samples cooled below c. 100°c in the early jurassic c2 episode (e.g. samples gc1077-62, -65, -66, -69). similarly, in the south-west, samples gc107758 and -59 cooled below 110°c in this episode while middle–upper jurassic vr sample (gc1077) afta sample (gc1077) afta sample (gc1104) lower cretaceous 5 km gc1077-103: basement >120°c, 325–250 ma (c0) 65–80°c, 115–45 ma (c4) gc1077-104: basement >100°c, >220 ma (c0) 80–90°c, 220–115 ma (c3) 45–70°c, 65–10 ma (c8) gc1077-102: basement >100°c, 350–220 ma (c0) 85–95°c, 235–35 ma (c4) 45–80°c, 30–0 ma (c8) gc1077-13: 130–125 ma 90–100°c, >50 ma (c4) 60–80°c, 40–0 ma (c8) gc1104-23: 140–136 ma 75–90°c, 130–10 ma (c4) 20–75°c, 20–0 ma (c8) gc1104-20: basement >105°c, 205–165 ma (c2) 75–85°c, 20–0 ma (c8) gc1104-22: basement >100°c, 245–190 ma (c0) 75–85°c, 135–15 ma (c4) 30–70°c, 15–0 ma (c7) gc1077-12: 130–125 ma 75–100°c, >20 ma (c4) 40–80°c, 40–0 ma (c8) gc1077-16: basement >105°c, 370–280 ma (c0) 75–95°c, 280–50 ma (c4) 40–70°c, 40–0 ma (c8) gc1077-101: 140–136 ma 80–100°c, 140–35 ma (c4) 35–75°c, 40–0 ma (c8) gc1077-83: basement >100°c, 210–150 ma (c2) 80–85°c, 150–45 ma (c4) 40–65°c, 20–0 ma (c8) gc1077-15: 161–156 ma 80–105°c, >80 ma (c4) 55–70°c, 45–0 ma (c8) 84°c,84°c,84°c 72°c,84°c,96°c 91°c,91°c,90°c 81°c,80°c 19°w 76 °1 5' n 76 °1 5' n 19°w fig. 15 thermal-history solutions derived from afta and vr data for store koldewey. stratigraphic ages listed alongside the sample numbers. palaeotemperature and timing constraints are listed for each sample together with the regional cooling episode assigned to the constraints. maximum palaeotemperatures derived from the mean vr value in each sample are compared with the thermal history solutions derived from afta data in samples from the same area. the maximum palaeotemperatures indicated by the vr data in mesozoic sediments are consistent with mid-cretaceous c4 palaeotemperatures derived from afta, showing that they cooled from maximum post-depositional palaeotemperatures in the c4 episode. map location in fig. 11. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 33 of 154 www.geusbul let in.org nearby samples cooled from similar palaeotemperatures in either the middle triassic c1 or early jurassic c2 episodes. given the sporadic evidence for this episode in the interior of the study area it appears likely that episode c3 affected the entire region but was overprinted in more coastal regions by later episodes. 4.5.5 mid-cretaceous episode (c4) cooling began between 95 and 90 ma. palaeotemperatures characterising episode c4 are recognised across much of the same area as the early jurassic c2 episode (c4 in fig. 14). palaeotemperatures are around 80 to 100°c across most of the study area. some locations, particularly at low elevations, give values in excess of 100°c. this episode is not recognised at coastal locations between the northern shores of hold with hope and jameson land, reflecting the dominance of cenozoic events at these locations. however, the relative uniformity of the episode, where it is revealed, suggests that it probably also affected those areas but has been overprinted by later events. 4.5.6 early eocene episode (c5) cooling began c. 55 ma. episode c5 is more restricted in extent than earlier episodes (c5 in fig. 14), focussed mainly around jameson land, where palaeotemperatures are around 100°c or above. five afta samples and two vr samples from northerly locations around clavering ø also show palaeotemperatures >100°c in this episode. these samples were collected close to mapped intrusive bodies of similar age. this, together with the sporadic occurrence of this episode, suggests that c5 is directly related to intrusive bodies. the abundance of early eocene intrusions across jameson land (noenygaard 1976) suggests a similar mechanism. we therefore interpret the effects of episode c5 to reflect local heating associated with the widespread eocene intrusive activity across the region (fig. 5; larsen et al. 2014). 4.5.7 end-eocene episode (c6) cooling began between 37 and 35 ma. episode c6 is observed at coastal locations south of kuhn ø (c6 in fig. 14). palaeotemperatures from samples at sea 74 °n 74 °n 21°w 21°w 74 °n 74 °n 21°w 21°w 74 °n 74 °n 21°w 21°w 74 °n 74 °n 21°w 21°w 74 °n 74 °n 21°w 21°w 74 °n 74 °n 21°w 21°w 74 °n 74 °n 21°w 21°w palaeotemp. (°c) >110 90–100 80–90 70–80 60–70 <60 100–110 palaeotemp. (°c) >110 90–100 80–90 70–80 60–70 <60 100–110 palaeotemp. (°c) >110 90–100 80–90 70–80 60–70 <60 100–110 palaeotemp. (°c) >110 90–100 80–90 70–80 60–70 <60 100–110 145–140 ma c3 95–90 ma c4 37–35 ma c6 c. 10 ma c8 cø gb hwh my vt4 vt14 vt14 vt4 vt3 vt2 vt1 vt10 vt9 vt16 vt17 vt3 hl gb kø sø wf vt16 vt16 vt9 vt10 vt1 vt3 vt2 vt4 d pl hl pd m f d f 25 km 25 km 25 km 25 km 74 °n 74 °n 21°w 21°w 74 °n 74 °n 21°w 21°w 74 °n 74 °n 21°w 21°w 74 °n 74 °n 21°w 21°w fig. 16 focussed view of palaeotemperatures from afta and vr around clavering ø for episodes c3, c4, c6 and c8 (table 1). cø: clavering ø. d: dombjerg. df: dombjerg fault. gb: giesecke bjerge. hl: hudson land. hwh: hold with hope. kø: kuhn ø. my: myggbukta. pdmf: post-devonian main fault. pl: payer land. sø: sabine ø. vt: vertical transect. wf: wollaton forland. sample locations and geology legend in fig. 11. map location in fig. 14. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 34 of 154 www.geusbul let in.org level suggest a progressive westerly decrease from the coast to inland, while this event is most emphatically expressed around the eocene intrusive centres in the eastern part of traill ø, in myggbukta and sabine ø. highest c6 palaeotemperatures in excess of 110°c are recognised from locations close to late eocene igneous intrusions on traill ø, while samples near sea level in wollaston forland and hold with hope reached c. 100°c at this time. a decrease of palaeotemperatures with increasing elevation is also evident in samples from vertical transects on traill ø northwards to hudson land (vt6, vt4; appendix 2.3). the absence of this event in samples away from intrusions indicates distinct hot spots around these intrusive centres. similar timing between episode c6 and the age of intrusive bodies at these three hot spots, provides further evidence that c6 is in some way related to these intrusions (section 5.7). 4.5.8 early miocene episode (c7) cooling began between 20 and 18 ma. palaeotemperatures for episode c7 are highly localised to northern jameson land (c7 in fig. 14). palaeotemperatures exceed 110°c around mestersvig and wegener halvø. values are slightly lower around schuchert dal (80– 100°c). it seems likely that palaeotemperatures in this episode are related to the intrusive centre of werner bjerge, located within the focus area of this episode. the intrusion age of the malmbjerget stock (25.8 ma; fig. 5) within the centre suggests that cooling was somewhat protracted. late-stage intrusive and hydrothermal activity are alternative possible explanations. the contrast between this episode affecting the southern shores of kong oscar fjord and the end-eocene c6 episode on the northern shore is a striking feature. the timing constraints from afta define this difference very clearly (appendix 1.6). early miocene palaeotemperatures on the southern shore probably overprinted the end-eocene signature in that region. cooling attributed to this time interval is identified in data from sample gc1077100 of the early miocene intrusion on hvalros ø (20.3 ± 0.2 ma; larsen et al. 2014) and interpreted as immediate post-emplacement cooling. 4.5.9 late miocene episode (c8) cooling began at c. 10 ma. of all the palaeothermal episodes recognised in this study, episode c8 is the most regional in extent, recognised in almost all samples (c8 in fig. 14). palaeotemperatures characterising this episode are generally around 60 to 80°c at inland locations close to sea level, decreasing with increasing elevation in vertical transects (appendix 2.3). note that the timing constraints from afta data attributed to this episode in individual samples show some inconsistencies, and in some cases, it remains possible that palaeotemperatures related to the c8 event may in fact relate to either the early miocene c7 or early pliocene c9 episodes. this problem relates partly to the relatively low palaeotemperatures characterising this episode in many samples, but also reflects the difficulty in resolving separate episodes that are closely spaced in time. a clear hotspot of palaeotemperatures around 100°c is evident at some locations around the eastern end of kong oscar fjord towards the eastern end of traill ø and near mestersvig, where afta data are dominated by the end-eocene and early miocene episodes, respectively. late miocene palaeotemperatures decrease in all directions onshore away from these areas, defining a broadly concentric pattern (see section 6.3). 4.5.10 early pliocene episode (c9) cooling began at c. 5 ma. episode c9 is recognised mainly around the eastern end of traill ø (c9 in fig. 14). this area is the focus of highest palaeotemperatures in the early and late miocene episodes (c7, c8), and the early pliocene episode is defined most clearly in samples for which late miocene palaeotemperatures around 90°c or above have been defined. in these samples, it is clear that two discrete cooling episodes are required within the last 10 myr or so, due to higher palaeotemperatures in the earlier episode. we infer that the early pliocene episode was most likely of regional extent and has affected the entire region, but it is only definitively recognised at these locations due to the higher temperatures that remained following late miocene cooling (see section 6.3). early pliocene cooling is also interpreted for one inland sample on payer land, for which the palaeotemperature defined from afta is much lower than those attributed to the late miocene episode at higher elevations (gc1077-32; vt10; appendix 2.3). since several samples from traill ø and wegener halvø define both the c8 and c9 episodes, it seems likely that both episodes did affect the whole region. but over most of the region, with late miocene palaeotemperatures <60°c in samples at elevations less than 1 km, resolution of distinct c8 and c9 episodes is not possible, and in most cases, the effects of both are merged into a single event attributed to the late miocene. 4.6 palaeogeothermal gradients inferred from afta in vertical transects 4.6.1 principles of the method where palaeotemperatures can be defined from afta over a range of elevations (in outcrop) or depths (in wells), the rate of increase with depth allows definition of the range of palaeogeothermal gradients that can explain the data. extrapolation of fitted palaeogeothermal gradients from the corresponding unconformity surface to an appropriate value of palaeosurface http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 35 of 154 www.geusbul let in.org temperature then allows estimation of the amount of additional section required to explain those palaeotemperatures (green & duddy 2012). estimating amounts of removed section in this way is subject to a number of critical assumptions. green et al. (2013) discussed these in detail, and we summarise the major points here. the most critical assumption is that the palaeogeothermal gradient was linear throughout the entire section at the time of maximum palaeotemperatures. this assumption may be invalid if non-burial-related heating (e.g. confined fluid flow) or major differences in lithology produce severely non-linear palaeotemperature profiles, which can result in linear extrapolation overestimating the amount of removed section. for this reason, the values quoted here should be viewed as upper limits to the total amount of section removed. selection of an appropriate surface temperature is also critical. another crucial assumption is that where multiple episodes of exhumation occur within a period not represented by preserved rock, these methods define the thickness of rock that was present at the onset of cooling in each episode. they do not define the total amount of rock removed in any single event (except for the most recent event) because the history between the separate cooling episodes is not constrained (see appendix 1). for example, in this study, the amount of reburial (if any) prior to the onset of late miocene cooling/exhumation is not known. therefore, it is not clear what proportion of the total amount of section present above the palaeogene basalts, when end-eocene cooling began, was removed prior to the onset of late miocene cooling. if sediments or lavas deposited between the two cooling episodes are still present, the situation is different. in this case, the underlying rock must have been exhumed to the surface and thus cooled to the surface temperature at the time of deposition. the palaeosurface temperature at the appropriate time can be used as a constraint on the reconstruction of the burial and exhumation history of the underlying rock. in addition, the sedimentary remnant provides clear evidence of the additional section (albeit seriously diminished) that is required to explain the palaeotemperatures reached in the subsequent episode. 4.6.2 palaeotemperature profiles in vertical transects we collected samples from 17 vertical transects (locations in fig. 11). in appendix 2 we present a detailed discussion of the results in each transect. for each transect we have avoided including samples over too wide a horizontal distance, to guard against lateral variation in the magnitude of palaeothermal effects, which might arise due to variation in heat flow, differential uplift (tilting) or offsets across major faults. comparing results from adjacent profiles suggests that this approach is justified (appendix 2.3). reliable definition of palaeotemperature profiles is only possible for the more recent palaeothermal episodes because most samples were totally annealed in the pre-cenozoic episodes, providing only minimum palaeotemperature estimates. only in rare situations do samples provide finite palaeotemperature constraints in earlier episodes over a sufficient range of elevations to allow determination of the allowed ranges of palaeogeothermal gradients and removed section. 4.6.3 estimates of palaeogeothermal gradients for a number of the vertical transects, sufficient constraints are available to justify quantitative determination of palaeogeothermal gradients and removed sections, and here we summarise the results presented in appendix 2. maximum likelihood values of palaeogeothermal gradient for the analysed transect, together with the corresponding 95% confidence limits, are compared in fig. 17. in most cases, the data only allow definition of broad constraints on the range of palaeogeothermal gradients. this is due partly to limited vertical ranges over which samples are 0 10 20 30 40 50 60 70 80 0 pa la eo ge ot he rm al g ra di en t ( °c /k m ) 10 ma (c8) 37–35 ma (c6) 100–90 ma (c4) 145–140 ma (c3) v t1 v t2 v t4 v t5 v t6 v t8 v t9 v t1 0 v t1 1 v t1 2 v t1 3 v t1 6 v t1 3 v t4 v t6 v t1 2 v t1 6 v t1 v t3 v t1 7 v t9 fig. 17 comparison of palaeogeothermal gradients in four palaeothermal episodes (c3, c4, c6, c8; table 1) determined from afta-derived palaeotemperatures in vertical transects (vts). maximum likelihood estimates (squares) with upper and lower 95% confidence intervals are shown. in most cases, only broad constraints are possible on the range of palaeogeothermal gradients allowed by the data. this is due partly to limited vertical ranges over which samples are available, and partly due to broad palaeotemperature constraints. the best-controlled values are highlighted in boxes, vt6, vt8 and vt10 for the late miocene c8 episode and vt4 for the end-eocene c6 episode. all three of the best-constrained estimates for the late miocene episode define maximum likelihood values around 25°c/km. while results from most of the other transects are consistent with this value, end-eocene constraints from vt4 and vt6 favour higher values that may be due to enhanced heat flow around intrusive bodies of the same age. location in fig. 11. further discussion in appendix 2. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 36 of 154 www.geusbul let in.org available and partly due to broad palaeotemperature constraints in most cases. only vt1 provided useful constraints on the earliest cretaceous c3 episode, suggesting a value close to 20°c/ km, while no constraints are possible on any of the earlier episodes in any of the transects. vt9 and vt16 provide broad constraints on the mid-cretaceous c4 episode, again suggesting values around 20°c/km. five transects provide constraints on the end-eocene c6 episode, of which the tightest constraint is from vt4, suggesting a value around 30°c/km. thirteen transects provide some constraint on the late miocene c8 episode, of which vt6, vt8 and vt10 are most tightly constrained. results from these three transects define maximum likelihood values around 25°c/km, while results from the other transects are consistent with this value. 4.6.4 summary in general, estimates of palaeogeothermal gradients in different events are consistent with values around 25°c/ km and rule out values in excess of c. 40°c/km (fig. 17). early jurassic palaeotemperature constraints from the vt3 transect (discussed in appendix 2) suggest a much higher palaeogeothermal gradient, but we regard this as an outlier, perhaps due to faulting, and it is not discussed further. on this basis, we favour an interpretation in which all pre-cenozoic palaeothermal episodes, as well as the late miocene and pliocene episodes, were characterised by palaeogeothermal gradients around 25°c/km. such gradients imply that heating reflects the effects of deeper burial prior to cooling due to exhumation. this is emphasised by the constant offset in palaeotemperature profiles characterising the c1, c3 and c8 episodes in vt1, clavering ø, while the c1, c2 and c8 episodes in vt17, north-east clavering ø (appendix 2.3) provide another example. a palaeogeothermal gradient around 25°c/km also applies to end-eocene c6 palaeotemperatures at locations away from intrusive bodies, while vt4 and vt6 define higher palaeogeothermal gradients suggesting locally enhanced heat flows near late eocene intrusions on traill ø and eastern hudson land. the more localised palaeotemperatures defining the early eocene c5 and early miocene c7 episodes are interpreted as localised effects of contact heating around intrusive bodies. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 37 of 154 www.geusbul let in.org in this chapter we compare the timing of the ten cooling episodes defined from afta (table 1) with the regional tectono-stratigraphic framework for north-east greenland (figs 4, 5), to provide insights into the nature of each episode. we also compare the afta results from northeast greenland with two thermal-history studies based on afta from east greenland between 61 and 70°n to establish which episodes of cooling and exhumation have affected the entire margin of east greenland. note that fig. 14 displays regional palaeotemperature maps for episodes c0 to c9, whereas fig. 16 displays detailed palaeotemperature maps for episodes c3, c4, c6 and c8 in the clavering ø region. 5.1 late carboniferous episode (c0) the c0 episode (beginning between 320 and 300 ma) correlates with the regional latest carboniferous to mid-permian unconformity (ranging from about 310 to 280 ma), which separates the upper permian sediments of the huledal formation from the underlying, tilted and truncated latest devonian to carboniferous strata of the traill ø group (fig. 18; surlyk 1990; vigran et al. 1999; stemmerik 2000). this unconformity marks the most profound change in tectonic style and overall depositional environment in the post-caledonian development of east greenland (surlyk 1990). this is supported by our results that demonstrate removal of a kilometre-thick rock column between the onset of c0 exhumation and the deposition of permian sediments on clavering ø, where a basement sample (gc1077-86) in depositional contact with upper permian sediments, cooled from >105°c to a late permian surface temperature of c. 25°c during episode c0. the unconformity is represented by a sub-permian peneplain covered by the upper permian sediments of the huledal formation (haller 1971; surlyk et al. 1984, larsen 1988; surlyk 1990, 2003). this, together with the recognition of this episode over a wide area (c0 in fig. 14) suggests that the c0 cooling episode reflects exhumation of present-day basement terrains, for example, west of jameson land and traill ø during the interval represented by this unconformity. this episode of regional exhumation resulted in the formation of the peneplain, which was subsequently buried below the mid-permian huledal formation. 5.2 middle triassic episode (c1) the c1 episode (beginning c. 240 ma) post-dates the late permian to lower triassic rift succession and correlates with the deposition of the thick siliciclastic wedges of the pingo dal group in the sedimentary basins on jameson land and traill ø (fig. 18). these units overly the marine mudstones of the wordie creek group on northern jameson land (clemmensen 1980; bjerager et al. 2006). extensive c1 cooling in the basement areas west of these basins (c1 in fig. 14) suggests that this event represents regional erosion of the areas to the west, providing sedimentary input to the basins (the pingo dal group and younger units of the scoresby land group). this indicates differential movement on the pdmf system at this time in the southern part of the study area. in the north, however, c1 exhumation dominates the area east of the pdmf system, suggesting a different tectonic history. the early–middle triassic rifting and regional exhumation represents an important phase during the break-up of pangaea (seidler et al. 2004; oftedal et al. 2005). 5.3 early jurassic episode (c2) the c2 episode (beginning c. 180 ma) corresponds to major regional uplift and erosion in east greenland around the early–middle jurassic boundary that has been established from stratigraphic and sedimentological observations (figs 18–20; surlyk & ineson 2003; surlyk 2003). we therefore interpret the c2 episode as representing regional denudation that affected areas west, east and north of the jameson land basin (c2 in fig. 14). the c2 episode could not be resolved in the afta data from jameson land, where both lower jurassic and triassic strata are present. the extensive denudation that started in the early jurassic eventually provided siliciclastic input to the jurassic basins; in particular corresponding to the middle jurassic sandstones of the charcot bugt and pelion formations that overlie triassic and upper permian sediments or peneplained basement on milne land and kuhn ø (see fig. 44 in surlyk 2003). the c2 episode slightly predates the middle jurassic (bajocian) sequence boundary at c. 170 ma in the jameson land basin but overlaps with the toarcian (183–174 ma; cohen et al. 2013) uplift and erosion on wollaston forland and kuhn ø indicated (fig. 20; surlyk & ineson 2003). given that the deposition of the pelion formation that rests on the jurassic peneplain on kuhn ø began at the bajocian–bathonian transition (c.170 ma), we can estimate the duration of exhumation during the c2 episode to be in the order of 10 myr. afta data provide evidence for early jurassic exhumation of the basement terrains in the interior of the study area. however, lower jurassic sediments are present in the jameson land basin, so whereas the basin subsided, the areas to the west of the basin were uplifted, and this implies that the pdmf system west of 5. integration of afta results with geological evidence http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 38 of 154 www.geusbul let in.org ma era stage/ age 0 10 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 210 220 230 240 250 260 270 280 290 300 310 320 330 340 350 paleocene eocene oligocene miocene pliocene pleistocene series/ epoch system/ period middle middle upper upper lower upper guadalupian lopingian lower lower middle mississippian upper mississippian lower penns. middle penns. upper penns. cisuralian lower mississippian messinian tortonian serravallian langhian aquitanian chattian rupelian priabonian bartonian lutetian thanetian selandian danian maastrichtian campanian santonian coniacian cenomanian albian aptian hauterivian valanginian berriasian tithonian kimmeridgian callovian bathonian bajocian aalenian pliensbachian sinemurian hettangian car nian ladinian anisian olenekian induan rhaetian norian roadian wordian capitanian wuchiapingian changhsingian kungurian artinskian sakmarian asselian gzhelian kasimovian moscovian bashkirian serpukhovian visean tournaisian oxfordian barremian toarcian turonian ypresian burdigalian pa le o zo ic m es o zo ic c en o zo ic c a r bo n if er o u s pe r m ia n tr ia ss ic ju r a ss ic c r et a c eo u s pa la eo g en e n eo g en e wollaston foreland – kuhn ø w e s n ? ? wollaston forland gp vardekløft gp bernbjerg fm foldvik creek gp store koldewey w e pelion fm traill ø – hold with hope s n w e traill ø gp foldvik creek huledal fm gp gråklint mb pingo dal gp gipsdalen gp fleming fjord gp vardekløft gp bernbjerg fm wollaston forland gp steensby bjerg fm home forland gp brorson halvø gp jackson ø gp månedal fm pelion fm jameson land basin s n ? ? w e kap stewart gp neill klinter gp traill ø gp foldvik creek wordie creek gp gp vardekløft gp pelion fm hesteelv fm raukelv fm hareelv fm sw ne milne land hartz fjeld fm kap leslie fm charcot bugt fm ups lps sub-permian peneplain sub-paleocene etch surface mid-jurassic etch surface c9 c8 c7 c6 c5 c4 c3 c2 c1 c0 fig. 18 caption and legend on the next page. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 39 of 154 www.geusbul let in.org jameson land was again active during the c2 episode. on kuhn ø, in the northern part of the study area, the presence of a peneplain below middle jurassic sediments indicates that uplift and erosion not only affected the hinterland but also parts of the rift. this agrees with the interpretation of surlyk (1978a) that early jurassic uplift and erosion was caused by a dome centred north of jameson land. however, our results show that the exhumation also affected areas far into the hinterland. the erosional products derived from c2 exhumation, for example the pelion formation, eventually reached the rift basins along most of the margin, and it thus seems likely that there were active movements also along the northern part of the pdmf system during this episode. 5.4 earliest cretaceous episode (c3) the c3 episode (beginning between 145 and 140 ma) coincides with the most important phase of extensional tectonics in the mesozoic that was initiated in middle volgian times (tithonian–berriasian transition, c. 145 ma; surlyk & ineson 2003). during this phase, the area north of kong oscar fjord was broken up into tilted blocks and deposition was almost exclusively coarse grained in proximity to active faults (surlyk 2003; figs 19, 20). the primary area where the c3 episode has been recognised overlaps with the area influenced by the latest jurassic to earliest cretaceous rift climax that led to the deposition of the deep-water, conglomerate-dominated wollaston forland group (c3 in figs 14, 16; surlyk 1978b, 2003; surlyk & korstgård 2013; henstra et al. 2016a). in this region, centered on clavering ø, the afta data define the c3 episode in samples west of the dombjerg fault (surlyk 1978b). this fault system constituted the western boundary fault during the rift climax and formed the high and steep, western coastline of the sea in the earliest cretaceous (surlyk 1978b). the source area for the wollaston forland group was the area to the west of the boundary fault with the coarsest basin fill derived from the crestal areas along the fault in agreement with outline of the region affected by exhumation during episode c3 (surlyk 1978b, 2003). we therefore interpret this episode as the result of regional uplift and erosion that accompanied the rift climax (fig. 18), not only along the margin of the rift but also of hinterland areas at considerable distance base-of-slope fans channel �lls marine mudstone marine carbonate, including chalk seas deep marine sandstone (beyond shelf break) shallow marine sandstone (shelf) arid terrestrial sedimentary environments lithological symbols other symbols anhydrite/gypsum coal layers sandstone siltstone mudstone/shale organic rich mudstone/shale limestone prograding shelf-margin sediment wedge volcanics extrusive igneous rocks mainly ma�c rifting/extension humid terrestrial/ marginal marine dolomite/evaporites/ limestone conglomerate bioherm ice-rafted dropstone intrusive magmatism tectonic and orogenic events inversion/compression/ uplift planation surface etch surface formation of peneplain time interval during which cooling began peneplains fig. 18 (continued) regional cooling episodes recognised from afta data in this study (table 1), time intervals for the formation of peneplains and the lithostratigraphic scheme for north-east greenland (fig. 4). vertical extent of the horizontal, coloured bars indicates uncertainty in cooling episode onset. exhumation in episodes c0–c4, c6, c8 and c9 were of regional extent whereas episodes c5 and c7 were local, partly related to magmatic activity. lps: lower planation surface. ups: upper planation surface. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 40 of 154 www.geusbul let in.org from the rift basins (e.g. sample gc1077-71, about 150 km west of the dombjerg fault). given the widespread, yet sporadic, occurrence of this event west of milne land and in the interior (west of ymer ø and clavering ø) it seems likely that exhumation during episode c3 affected a much wider region. this assertion is supported by the contemporaneous siliciclastic deposits of the hartz fjeld formation (milne land) and raukelv formation (jameson land). in the interior, it also seems likely that some degree of mesozoic faulting has taken place because samples with high c3 palaeotemperatures are adjacent to samples with high c2 or c1 palaeotemperatures. on the north-east greenland shelf, the base of deposits of inferred latest jurassic – early cretaceous age defines a characteristic low-angular, erosional truncation of the underlying strata (hamann et al. 2005). hamann et al. (2005) correlated this unconformity to the mid-volgian unconformity (c. 145 ma) at the base of the wollaston forland group onshore and interpreted it to mark the onset of rifting in the latest jurassic. we suggest that this offshore unconformity corresponds to exhumation onshore. 5.5 mid-cretaceous episode (c4) the c4 episode (beginning between 95 and 90 ma) most likely corresponds to the mid-turonian, erosional base of the jackson ø group that records a phase of basin reorganization (fig. 18; bjerager et al. 2020). the exhumation involved erosion of jurassic and triassic wollaston forland hochstetter forland kuhn ø n s bernbjerg fm payer dal fm pelion fm bastians dal fm caledonian basement permian ugpik ravine mb ? ts muslingebjerg fm o�shore-marine mudstones transition-zone heterolith shallow-marine sandstones non-marine (�uvial) sandstones/mudstones coal c1 c. 240 ma c2 c. 180 ma c3 145–140 ma c4 95–90 ma c0 320–300 ma fig. 19 jurassic stratigraphy in the wollaston forland area along a n–s profile and timing of the pre-cenozoic cooling episodes estimated from afta (table 1). the erosional unconformity at the base of the jurassic is caused by early jurassic doming and regional uplift of northern east greenland (surlyk 1978a, 2003). earlier events may have contributed to the pre-jurassic denudation, e.g. during the middle triassic. the middle triassic c1 episode dominates the afta data to such a degree that the early jurassic c2 episode cannot be resolved along the profile. maximum palaeotemperatures (and burial) of the mesozoic sediments occurred prior to regional uplift and erosion during the mid-cretaceous c4 episode. samples from wollaston forland cooled below 100°c in the c4 episode, so no earlier events can be resolved from this area. solid line: cooling episode resolved along the transect. dashed line: cooling episode not resolved along the transect. ts: transgressive surface. modified after surlyk (2003). http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 41 of 154 www.geusbul let in.org sandstones, permian carbonates and a wide spectrum of caledonian-derived crystalline clasts that were reworked into the månedal formation (surlyk & noenygaard 2001). note that upper cretaceous strata of post-turonian age are only known from the easternmost parts of wollaston forland and hold with hope and locally on traill ø and geographical society ø (donovan 1957; parsons et al. 2017; nøhr-hansen et al. 2019; bjerager et al. 2020). however, the relative uniformity of the cooling in episode c4 across the entire study area outside the jameson land basin and the coastal areas, suggests that exhumation affected most of the study area – even the coastal areas where the afta data were overprinted by later events (c4 in fig. 14,16). afta data do not record cooling related to the pronounced mid-albian (c. 105 ma) fault-associated unconformity that is observed on hold with hope, clavering ø and wollaston forland (kelly et al. 1998; whitham et al. 1999; bjerager et al. 2020). this indicates that the fault activity predominantly occurred during subsidence of hanging-wall blocks. 5.6 early eocene episode (c5) the sporadic occurrence of cooling (beginning at c. 55 ma), together with the clear proximity of these samples to mapped intrusive rocks of similar age, suggest that episode c5 is related to intrusive bodies (fig. 21). this is clearly the case on jameson land, where the extent of this episode overlaps with intensive palaeogene intrusive activity, and where a dyke is dated to c. 53 ma (noenygaard 1976; larsen et al. 2014). except for the few isolated occurrences in the north, episode c5 is strikingly focussed within the vicinity of jameson land. we therefore suggest that the palaeotemperatures characterising this episode are most likely due to either contact or hydrothermal effects associated with igneous activity. no evidence for any regional maastrichtian–paleocene or paleocene–eocene exhumation has been identified in this study, but there is a hiatus at the base of the paleocene succession and material derived from reworked upper cretaceous units are present within the paleocene deposits (nøhr-hansen et al. 2011). as discussed in section 2.3, thick volumes of volcanic and sedimentary rocks accumulated across shelf drowning shelf drowning marine inundation overall transgressive trend minor uplift basin reorganisation gentle block rotation, initial sedimentary progradation regional backstepping deepening – persistent regional subsidence (block rotation) progradation of shelf systems uplift and incision infill/drape of erosional relief basin segmentation and block rotation uplift/erosion onlap onto pre-jurassic basement, sedimentary progradation gentle block rotation regional backstepping syn-rift phase pre-rift phase uplift fault-controlled subsidence (minor/major) east greenland jameson land – milne land wollaston forland – kuhn ø ?uplift rift climax rift climax regional thermal subsidence chronostratigraphyma ryazanian volgian berriasian tithonian kimmeridgian oxfordian callovian bathonian bajocian aalenian toarcian pliensbachian sinemurian hettangian rhaetian c re ta ce ou s ju ra ss ic tr ia ss ic lo w er repp u re wol m id dl e u pp er u m l u l u m l u m l u m l u l u m u l m u l u l 210 200 190 180 170 160 150 140 c2 c3 fig. 20 summary of the main tectonic events and trends in jurassic – early cretaceous basin evolution in east greenland compared with the timing for the onset of cooling in the early jurassic c2 and the earliest cretaceous c3 episodes. note that c2 is not resolved on jameson land as indicated by the dashed green band. arrows indicating uplift from surlyk & ineson (2003). modified from surlyk & ineson (2003). http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 42 of 154 www.geusbul let in.org the east greenland margin in response to subsidence during and after break-up. therefore, maximum palaeotemperatures were likely achieved after, rather than before, continental break-up in the earliest eocene, and consequently any early palaeogene exhumation cannot not be detected by afta. 5.7 end-eocene episode (c6) the c6 episode (beginning between 37 and 35 ma) correlates with the regional post-breakup unconformity that separates the palaeogene basalts from quaternary deposits. the episode is recognised in afta data in samples of palaeogene and older age in a band from wollaston forland over ymer ø to traill ø and in liverpool land, south-eastern jameson land and blosseville kyst (c6 in figs 14, 16). evidence of this episode is most strongly expressed on eastern traill ø, but absent on western and northern jameson land – possibly due to the local effects of the early eocene c5 and the early miocene c7 episodes. constraints on the c6 episode are tight along some vertical transects. for example, on hudson land (vt4) where the palaeogeothermal gradient varies between 20 and 40°c/km (appendix 2.3). additional burial is required to explain the palaeotemperatures in this episode, with an additional contribution from elevated heat flow around intrusive bodies. the summit intercept at vt4 for this episode is about 70°c, which requires that 2 km or more have been removed since the late eocene at this location. palaeogene basalts are not present on the summits of hudson land, but they do occur on nearby summits, for example on clavering ø. it thus seems likely that the end-eocene palaeotemperatures were c9 c8 c7 c7 c7b c7b c7a c7a c6 10 0 ma ma 20 30 40 50 60 63°n 64°n 65°n drift rift 66°n 67°n 68°n 69°n 70°n 71°n 72°n 73°n 74°n 75°n vindtop fm malmbjerget hvalros ø borgtinderne bjørn flammefjeld wiedemann fjord kangerlussuaq granite ejnar mikkelsen nephelinite diatreme ikâsangmit nûk myggbukta, late sheet werner bjerge kap simpson kap parry kap broer ruys (myggbukta, inferred) ørsted dal mols bjerge shannon nunataks freycinet bjerg svinhufvud bjerge igtertivâ fm above cgl. igtertivâ fm below cgl. clasts in conglomerate pr. wales bj. fm gardiner granitekærven skaergaard sulugssut dykes imilik ii kap edvard holm nordre aputitêq igtutarajik pâtûlâjivit kruuse fjord kap gustav holm hole 988a hole seg01 hole 989b hole 990a lower series middle series hole 918d hole seg58 hole 917a imilik iii søndre aputitêq snout series lilloise kap boswell kræmer ø/ kangerlussuaq syenite gabbro dyke sill central intrusive complex lava succession plateau basalts lower basalts pa le oc en e o d p le gs 15 2+ 16 3 o d p le g 16 3x o d p le g 16 3 o d p le g 16 3x ka p g us ta v h ol m re gi on ki al in eq re gi on n ua lik re gi on ka ng er lu ss ua q sc or es by s un d ja m es on l an d tr ai ll ø g eo gr . s oc ie ty ø bo nt ek oe ø h ol d w it h h op e sh an no n w ol la st on fo rl an d bl os se vi lle ky st eo ce ne o lig oc en e m io ce ne p q mols bjerge sk jo ld un ge n borgtinderne c5 fig. 21 north–south extent of cenozoic cooling events identified from afta data along the east greenland margin (green et al. 2014; japsen et al. 2014; this study) compared to igneous activity from fig. 5 (larsen et al. 2014). coloured bands along the south–north axis represent the geographical extent of each episode. vertical extent of the coloured bands indicate uncertainty for the onset of the cooling episode. dashed bands represent areas where the presence of these episodes is inferred. the presence of the c9 episode between 68 and 70°n was inferred by japsen et al. (2014) by integration of afta data with results from stratigraphic landscape analysis (bonow et al. 2014). the c6, c8 and c9 episodes are interpreted as representing regional events of uplift and erosion in the end-eocene, late miocene and early pliocene episodes, respectively. the early eocene c5 episode is related to the intense magmatic activity during breakup. c7, c7a and c7b are discrete events along the coast that are partly related to thermal activity and partly to exhumation in the early miocene, late oligocene and middle miocene episodes, respectively. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 43 of 154 www.geusbul let in.org mainly due to burial below a cover of eocene basalts and sediments. see further discussion in section 6.1.1. we have not recognised episode c6 in the western part of our study area where the mid-cretaceous c4 episode dominates and thus apparently makes it difficult to resolve the effects of the end-eocene episode in the afta data. the effects of the end-eocene episode appear to decrease rapidly westwards, such that the boundary between regions showing the c4 and c6 episodes is almost mutually exclusive (fig. 14), although some degree of overlap does exist in the region around hold with hope to wollaston forland. see section 6.1.2 for a discussion on extent of exhumation during the c6 episode in relation to stratigraphic landscape analysis. in the central part of the region, this and later periods of uplift are represented by a north–south-trending, left-lateral fault system that crosscut the margin and reactivated older fault systems in a strike-slip tectonic regime, with south-east–north-west-oriented, maximum horizontal stress (guarnieri 2015). the c6 episode affected the entire east greenland margin (fig. 21). consequently, we interpret it as representing regional uplift resulting in kilometre-scale exhumation. results from traill ø suggest a significantly higher basal heat flow in that area. this is consistent with the presence of major igneous bodies of similar age in this region (price et al. 1997; larsen et al. 2014). off north-east greenland, steep prograding clinoforms occur above a horizon defining erosional incision that petersen (2019) dated broadly as late eocene to mid-miocene age (fig. 6e). the clinoforms are caused by extensive mass wasting into the outer part of the shelf due to tectonic uplift of the inner part of the shelf and onshore areas, south of store koldewey (petersen 2019). we suggest that the mass wasting is due to the uplift of the margin of north-east greenland during the end-eocene c6 episode, and thus that the horizon defining the erosional incision can be dated to about 35 ma. 5.8 early miocene episode (c7) the c7 episode (beginning between 20 and 18 ma) is restricted to the northern parts of jameson land around the igneous centre of werner bjerge and malmbjerget (c7 in fig. 14), suggesting that it does not reflect regional exhumation. although these centres are older (oligocene, fig. 21) than the onset of cooling in this episode, the spatial association of cooling with the intrusive centre implies a direct link. we therefore conclude that this episode represents a local palaeothermal anomaly associated with emplacement of the igneous centres on northern jameson land. however, a component of uplift and erosion is likely to be associated with episode c7 (see the discussion of landscape analysis of stauning alper in section 6.4). 5.9 late miocene episode (c8) the c8 episode (beginning at c. 10 ma) also correlates with the regional post-breakup unconformity that separates the palaeogene basalts from quaternary deposits (fig. 18) and is recognised across almost the entire study area (c8 in figs 14, 16). at most locations, palaeogeothermal gradients around 25°c/km provide a satisfactory description of the palaeotemperatures in this episode. we therefore interpret palaeotemperatures characterising this episode in terms of depth of burial, with cooling due to exhumation. while episode c8 is not recognised in the western parts of jameson land, because of the strong signature of the early miocene c7 episode there, it seems likely that late miocene cooling also affected that region to some degree. in the interior highlands, late miocene palaeotemperatures in samples that also record the mid-cretaceous c4 episode can be explained by palaeoburial below a late miocene palaeosurface, corresponding to the general summit level (ups) in these mountains (fig. 22). the c8 episode in the interior highlands can thus be explained in terms of the incision of the present-day relief below the present-day summits (ups), with little or no regional denudation. in contrast, over much of the coastal region, an additional section above the present-day summit levels must have been present when cooling began at c. 10 ma in this episode (section 6.2). this is demonstrated by c8 palaeotemperatures that extrapolate to likely values of palaeosurface temperatures at elevations well above the present-day topography (appendix 2.3; e.g. vt1–6, 9, 12–16). as palaeogene basalts are present in the summits at several localities, much of the additional kilometre or more of removed section must represent additional basalts and post-basalt sediments. results from the vertical transects in the coastal region typically require a cover above the summits of about 1 km in the late miocene (palaeogeothermal gradient 25°c/km, late miocene surface temperature 15°c; appendix 2). around eastern traill ø, c8 palaeotemperatures are in excess of 100°c close to igneous intrusions, and palaeogeothermal gradients are higher than in areas immediately to the west. however, similar amounts of missing section are inferred here as in the areas away from the intrusions. we therefore interpret late miocene palaeotemperatures over the study region as representing deeper burial, combined with an elevated heat flow in the vicinity of traill ø. this represents a diminishing signature of late eocene intrusive activity away from the main intrusive centres. this is probably also true in samples around the early miocene intrusions at werner bjerge. late miocene cooling therefore reflects regional exhumation, where cooling of inland samples reflects incision of the present-day relief below the present-day summit http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 44 of 154 www.geusbul let in.org level (ups). this cooling episode is therefore interpreted as representing the main uplift of the continental margin after break-up in the north-east atlantic. off north-east greenland, massive shelf progradation occurred above a distinct angular unconformity, the imu (fig. 6; hamann et al. 2005; berger & jokat 2008; døssing et al. 2016; petersen 2019). døssing et al. (2016) estimated the age of the imu to be middle to late miocene (15–10 ma) based on odp data (myhre et al. 1995; hull et al. 1996). the imu marks the termination of synrift deposition in the deep-sea basins and the accelerated widening of the fram strait, and it is linked to the onset of late miocene uplift and massive shelf progradation on the north-east greenland margin (døssing et al. 2016). we therefore suggest that the uplift onshore and the subsequent progradation offshore began during the late miocene episode, and thus that the imu can be dated to about 10 ma. 5.10 the early pliocene episode (c9) the c9 episode (beginning at c. 5 ma) is recognised only in areas where palaeotemperatures in the early or late miocene episodes (c7, c8) are particularly high (e.g. eastern traill ø; compare c7–c9 in fig. 14). this episode also correlates with the regional post-breakup unconformity that separates the palaeogene basalts from quaternary deposits. since the extent of the exhumation during this event cannot be established from afta or the stratigraphic record, we return to this in section 6.5 where we discuss episode c9 alongside evidence from stratigraphic landscape analysis. 5.11 comparison with results from adjacent areas of east greenland thermal-history studies based on afta data have previously been published for two regions of east greenland to the south of the present study region: south-east greenland (green et al. 2014; 61–66°n) and southern east greenland (japsen et al. 2014; primarily 68–70°n). compilation of the results of these studies and those of the present study show that the timing of cooling in many of the palaeothermal episodes recognised in each of these regions show a high level of consistency (fig. 23). the results presented here concerning the cenozoic development of north-east greenland reveal cooling in the end-eocene c6 and late miocene c8 episodes, representing exhumation, interpreted to have affected most (perhaps all) of the region (see sections 6.1 and 6.2). early pliocene c9 exhumation, resolved in restricted locations, is also inferred to have affected much of the region. we regard these three episodes as representing key phases of uplift in the regional development of the present-day continental margin of north-east greenland. end-eocene c6 and the late miocene c8 episodes correlate closely with events in southern east greenland and south-east greenland (fig. 23). in southern east greenland (68–71°n), episodes c6 and c8 represent the main events involved in the uplift of the present-day mountains of the continental margin (japsen et al. 2014). integration of afta data with results from stratigraphic landscape analysis (bonow et al. 2014), led japsen et al. (2014) to conclude that a significant episode of early pliocene uplift also affected that region. thus, our results define a consistent style of development of the continental margin in these two areas of east greenland. in south-east greenland (61–66°n), while cooling general summit level el ev at io n (k m ) 0.5 1.0 0.0 1.5 palaeotemperature (°c) 20 40 60 80 100 120 140 160 1800 2.0 c. 10 ma c8 25°c/km gc1077-65 gc1104-14 gc1077-62 store koldewey and inland areas stauning alper milne land inland south of 74°n fig. 22 late miocene c8 palaeotemperature constraints from afta in samples from four sub-regions in the interior basement terrains plotted against elevation. outlying samples in which the mid-cretaceous c4 episode is not identified are omitted. results from these regions are consistent with a linear gradient of 25°c/ km and extrapolate to a typical palaeosurface temperature at the general summit level of 2.0 to 2.5 km a.s.l. (corresponding to the ups, the upper planation surface). cooling from these palaeotemperatures over the last 10 ma can be explained by incision below the regional summit level. sample gc1104-14 for which episode c4 is not resolved, is included to illustrate the likely influence of unresolved cooling events. the results for this sample and for gc1077-65 from about the same elevation and area provide a clear example of the importance of unresolved multiple cooling episodes. only the constraints for sample gc1077-62 plot above the regional trend. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 45 of 154 www.geusbul let in.org in the 40–30 ma interval correlates closely with events in more northerly regions, only a single event is defined within the last 10 myr. nevertheless, based on the overall consistency a similar development seems likely in that region too (green et al. 2014). the end-eocene episode identified in this study is closely related to igneous activity around traill ø and other areas, although this event is dominated by exhumation over most of the region. many of the other cenozoic cooling episodes identified in this and previous studies also show close association with intrusive activity along much of the margin (fig. 21). early miocene c7 cooling identified in this study is interpreted as representing local effects related to intrusive bodies of similar age. in southern east greenland, japsen et al. (2014) also identified cenozoic cooling episodes of very restricted areal extent along blosseville kyst. a late oligocene episode was interpreted as reflecting localised uplift and erosion north-east of a major fault along kangerlussuaq. a middle miocene episode, coeval with the extrusion of the lavas of the vindtop formation, was interpreted as reflecting circulation of hot fluids, probably during some degree of crustal extension. these events have no counterparts outside those areas. early eocene cooling in north-east greenland and southern east greenland is also associated mainly with minor intrusions. however, from the southernmost tip of east greenland, green et al. (2014) reported evidence of a late cretaceous – early cenozoic episode, which was thought to represent exhumation due to major movements along east–west-oriented faults. thus, a somewhat more complex evolution is possible in that area compared to regions to the north. the mesozoic c1, c2 and c4 episodes identified in this study (middle triassic, early jurassic and mid-cretaceous, respectively) have correlative cooling events in other areas (fig. 23) and appear to have affected most of the margin of east greenland south of 78°n. in contrast, timing constraints on the earliest cretaceous episode in this study (145–140 ma) and south-east greenland (135–130 ma) are slightly different between the two regions. given the consistency of different cooling episodes in the three u ps lp s break-up e. pliocene (c. 5) l. miocene (c. 10) q north-east greenland 70–78°n pl mio oli eo pal ku kl j tr pe ca end-eocene (37–35) e. jurassic (c. 180) m. triassic (c. 240) e. cretaceous (145–140) m. cretaceous (95–90) c1 c2 c6 c8 c9 c3 c0 c4 l. carboniferous (320–300) ti m e (m a) e. eocene (c. 55)c5 lp s break-up l. miocene (c. 10) q southern east greenland 68–70°n pl mio oli eo pal ku kl j tr pe ca l. eocene (40–35) break-up l. miocene – recent (10–0) 30 0 20 0 10 0 65 60 40 20 0 30 0 20 0 10 0 65 60 40 20 0 30 0 20 0 10 0 65 60 40 20 0q south-east greenland 61–66°n pl mio oli eo pal ku kl j tr pe ca eocene–oligocene (40–30) e. jurassic (c. 175) e. triassic (240–235) e. cretaceous (135–130) mid-late cretaceous (100–85) e. jurassic (200–175) triassic (260–210) cretaceous (125–85) l. dev. – carb. (370–310) ’e. pliocene’ (c. 5) m. miocene (17–12) l. oligocene (27–25) e. eocene (55–50) u ps l. cretaceous – e. cenozoic (70–50) e. miocene (20–18)c7 formation of peneplain uplift event from stratigraphic landscape analysis time interval during which cooling began onset of cooling (ma) from afta(10–0) fig. 23 timing of post-devonian cooling episodes in east greenland estimated from afta data and time intervals for formation of peneplains. south-east greenland (green et al. 2014; 47 samples, 61–66°n); southern east greenland (japsen et al. 2014; 90 samples, primarily 68–70°n but samples as far south as 64°n), north-east greenland (this study, table 1; 217 outcrop samples, 70–78°n). the vertical extent of the horizontal, coloured bars indicates the uncertainty for the onset of the cooling episode. ca/ carb: carboniferous. dev: devonian. e: early or earliest. eo: eocene. j: jurassic. kl: lower cretaceous. ku: upper cretaceous. l: late. lps: lower planation surface. m: mid. mio: miocene. oli: oligocene. pal: paleocene. pe: permian. pl: pliocene. q: quaternary. tr: triassic. ups: upper planation surface. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 46 of 154 www.geusbul let in.org regions in fig. 23, it is possible that the difference in the earliest cretaceous constraints simply reflects minor statistical inconsistencies within the dataset. alternatively, some slight diachroneity across the region is possible. nevertheless, it is clear that early cretaceous cooling occurred at broadly similar times across the margin of east greenland from 61°n to 78°n, suggesting control by regional tectonic processes. the late carboniferous c0 cooling episode identified in this study coincides with a cooling episode in southern east greenland, again suggesting regional exhumation associated with tectonism. this episode is not recognised in south-east greenland, largely because later events dominate the history of that region where palaeozoic events were overprinted. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 47 of 154 www.geusbul let in.org subaerial palaeogene basalts occur onshore in east greenland between 68 and 75°n (henriksen et al. 2009). the basalts reach elevations up to 3.7 km a.s.l. (gunnbjørn fjeld; c. 68°n). offshore, between 63 and 76°n, correlative, subaerial basalts are truncated below the seabed and dip away from the land areas, until they are buried below more than 2 km of sedimentary cover farther offshore (fig. 24; berger & jokat 2008). in contrast to this present-day configuration, the east greenland margin was characterised by net subsidence during the eruption of the basalts (larsen 1990; pedersen et al. 1997; brooks 2011). early studies concluded that onshore uplift post-dated volcanism by about 20–30 ma, in agreement with more recent results (larsen & marcussen 1992; japsen et al. 2014). here we investigate the evidence for eocene subsidence and burial along the margin and for the present topography being younger than 10 ma. finally, we integrate the results from afta 100 km jameson land scoresby sund wollaston forland dombjerg kap brewster renland liverpool land gåseland a b c d e milne land continent–ocean transition palaeogene basalts onshore/o�shore palaeogene basalts at seabed palaeogene intrusion onshore/o�shore carboniferous–cretaceous sediments devonian sediments pre-devonian rocks fault basalts outlier 0 1 –1 2 –2 3 –3 4 el ev at io n (k m a .s .l. ) –4 upsups lps(lps) dombjerg cover at 35 ma top basalt wollaston forlandnunataks glaciers ? ? 50 km d at a ga p a b c d e maximum altitude in corridor topography along pro�le palaeogene basaltpalaeogene cover (inferred) projected surface elevation at 35 ma intra-miocene unconformity (imu) 72 °n 74 °n 76 °n 20°n 15°n25°n30°n 50 km fig. 24 geology of the northern part of the study area. upper panel: present-day extent of subaerial palaeogene basalts across the northern part of the study area, onshore (fig. 11) and offshore (henriksen et al. 2009). lower panel: onshore–offshore profile illustrating the change in elevation/depth of the subaerial palaeogene basalts from 2 to –4 km a.s.l. (a–e in upper panel). these indications of strong, differential vertical movements along the margin in post-basalt time were noted by larsen (1990). the late eocene projected surface elevation (pse) over dombjerg is based on extrapolation of palaeothermal constraints from samples on wollaston forland and eastern clavering ø (figs 25, 26). the topography along the inland part of the profile (i.e. ab) is not indicated because it is primarily defined by the ice surface. nunataks with a cover of basalts at ‘a’ are cut by the upper planation surface (ups; bonow & japsen 2021, this volume). the elevation of the lower planation surface (lps) at some distance from the profile is marked. the offshore unit above the intra-miocene unconformity (imu, 15–10 ma; døssing et al. 2016) is characterised by prograding successions that are truncated below the seabed along the coast (see fig. 6). offshore profile from berger & jokat (2008). 6. development of the continental margin of north-east greenland after break-up http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 48 of 154 www.geusbul let in.org with those from stratigraphic landscape analysis to produce a model that explains key features of the modern topography of north-east greenland. 6.1 eocene burial and exhumation 6.1.1 evidence for eocene subsidence/burial and uplift/exhumation the end-eocene projected surface elevation (pse) is a measure of the elevation of the late eocene land surface relative to present-day sea level. pse is calculated by adding the thickness of the additional cover required to explain the end-eocene palaeotemperature defined from afta in an outcrop sample to its present-day elevation (fig. 25). if the end-eocene surface was close to sea level at the onset of exhumation, the pse thus provides a measure of the total amount of rock uplift since the onset of exhumation relative to present sea level (japsen et al. 2014). the thickness of the additional cover is calculated from the end-eocene palaeotemperature, assuming a surface temperature of 20°c and a palaeogeothermal gradient of 25°c/km in the late eocene (30°c/km for vt3, vt4, 40°c/km for vt6, vt14). afta and vr data from some samples along the northern coast indicate heating to end-eocene palaeotemperatures in excess of 100°c (e.g. north-east wollaston forland). these values stand out as anomalies that are likely to be related to intrusive bodies similar to those at the eastern part of traill ø (fig. 26). the pse for the end-eocene c6 episode is fairly uniform at around 3 km (a.s.l.) over much of the region where the episode is resolved, possibly increasing to 3.5 km or more at some locations nearer the coast (fig. 26). the regional, end-eocene pse of c. 3 km is higher than all summits in the study area and thus also above the regional peneplains, ups and lps (chapter 3). this indicates that rocks now exposed in the summits from wollaston forland to liverpool land, were buried below a considerable thickness of rock when end-eocene exhumation began, typically by around 1.5 km or more. therefore, the ups, which is younger than the palaeogene basalts, was formed as a result of uplift and erosion beginning end-eocene. around clavering ø palaeogene basalts are exposed on many summits, and consequently these summits must have been buried below palaeogene basalts and eocene post-basalt sediments prior to the onset of uplift and erosion in the late eocene (fig. 24). erosion removed these deposits, leaving the remnant basalt outcrops as the only visible evidence that this additional section once existed. this is further underpinned by afta data from one sample of paleocene sandstone from hold with hope that reached a palaeotemperature of 65–85°c prior to end-eocene cooling, corresponding to burial below a significant paleocene–eocene cover (sample gc1077-40; c6 in fig. 16). geological evidence shows that the region south of scoresby sund was characterised by net subsidence during and after the eruption of the basalts at the paleocene–eocene transition (section 2.2). furthermore, palaeotemperatures estimated for samples of early eocene sandstones from blosseville kyst (fig. 27) indicate that a succession of eocene sediments about 2 km thick covered the region prior to the onset of end-eocene c6 cooling. consequently, subsidence and burial dominated along the margin of north-east greenland throughout the eocene. such a history of burial during post-rift thermal subsidence is to be expected from classic theory of continental stretching (mckenzie 1978; white & mckenzie 1988). pse = 2.8–3.0 km 80°–90°c common pse for all samples 0.2 km 1.0 km 3.0 2.6 2.8 3.4 0 1 2 el ev at io n (k m a . s . l .) 3 4 2. 4– 2. 6 km c ov er 1. 8– 2. 4 km c ov er 1. 0 –1 .2 k m co ve r 65°–80°c cover prior to exhumation present relief fig. 25 projected surface elevation (pse) corresponding to palaeo temperatures for a palaeothermal episode identified in afta data for samples within a minor area (vertical transect). the temperature range for each sample can be converted to estimates of thickness of the cover above each sample at that time, e.g. assuming a surface temperature of 20°c and a palaeogeothermal gradient of 25°c/km. pse at that time for each sample is the sum of the sample elevation above present sea level and the thickness of the former cover (since removed). pse is a proxy for uplift since the onset of exhumation if the surface at that time was near sea level. the common pse for all samples is the range of pse values that falls within the pse intervals for all samples. assuming a common palaeogeothermal gradient in different regions implies that areas characterised by an elevated palaeogeothermal gradient will have higher pses. such a situation may arise if the cover rocks removed during an episode of exhumation were relatively unconsolidated sediments with lower thermal conductivities compared to the underlying basement rocks (appendix 2). grey: present-day terrain. yellow: removed cover that was present prior to exhumation. orange: common pse for both samples. modified after japsen et al. (2014). http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 49 of 154 www.geusbul let in.org 6.1.2 extent of eocene cover and of endeocene exhumation the end-eocene c6 episode is resolved in the afta data in coastal areas in a band from blosseville kyst over liverpool land and jameson land to traill ø, hudson land, hold with hope, eastern clavering ø to wollaston forland (fig. 14). the event is, however, also identified in afta data from samples farther from the coast, for example west of the pdmf on western traill ø and eastern ymer ø. the region underwent post-rift burial throughout the eocene. given the relative uniformity of the palaeotemperatures across the region, it is likely that this cover was continuous between the locations where end-eocene palaeotemperatures are recognised from afta. as the end-eocene palaeotemperatures in the clavering ø region reflect post-rift burial below at least 2 km of palaeogene basalts and eocene sediments, it is likely that this cover extended further inland beyond where the presence of the cover can be constrained. 20°w30°w 30°w 20°w 76 °n 74 °n 72 °n 70 °n 70 °n 72 °n 74 °n 76 °n 3.1 hotspot 2.7 3.1–3.3 3.0–3.2 2.9 2.5–3.2 3.5–3.7 3.1 3.3 3.5–3.8 2.5 end-eocene projected surface elevation (km a.s.l.) 3.5–3.9 3.0–3.4 2.5–2.9 3.4–3.8 myggbukta wollaston forland clavering ø hudson land ymer ø hold with hope traill ø >3.7 2.9 hotspot hotspot 50 km vertical transect sample 1 3 4 4a 6 14 12 13 15 pro�le in fig. 40 blokelv-1 scoresby sund pdmf reactivated in cenozoic time borehole fig. 26 common projected surface elevation (pse; fig. 25) above present sea level for the end-eocene c6 episode for groups of samples outlined by thin black lines. pse is calculated for a surface temperature of 20°c and a palaeogeothermal gradient of 25°c/km (vt3, vt4 30°c/km; vt6 40°c/km). values for the intrusive centres are omitted. due to the uncertainty in elevation of the samples for vt7, data from the gc103, gc159 and gc202 reports were not used. at all sample locations, the end-eocene pse is higher than the present-day summits indicating that all summits, fjords and valleys were buried below a column of rocks at that time. pse increases over a short distance, e.g. along the northern coast of ymer ø and on hudson land between vt3 and vt4 (fig. 29). pdmf: post-devonian main fault. geology in fig. 3. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 50 of 154 www.geusbul let in.org the eocene cover most likely extended to the nunataks in the interior where lower eocene basalts are preserved (near letter a in fig. 24), and where the ups corresponds to the summits at c. 2 km a.s.l. consequently, much of the rock column removed during the formation of the ups across this central region was eocene basalts and sediments. note that the basalts on the nunataks are distinctly younger than those in the coastal region (fig. 5). this may indicate that younger basalts were part of a continuous cover extending to the coastal zone (where only a deeper level of basalt is preserved) and offshore. north of wollaston forland, the end-eocene cooling event cannot be resolved in the afta data. this is for instance the case for the samples from dombjerg and kuhn ø (vt16, vt9), located a few tens of kilometres from samples on wollaston forland and eastern clavering ø indicative of an end-eocene pse of 2.7 and 3.4 km a.s.l., respectively (fig. 26). since the lps on clavering ø is 1.2 km a.s.l., the palaeogene basalts now exposed on the lps must have been buried below an eocene cover of about 1.5 km in the late eocene. if we consider that the eocene burial reflects post-rift thermal subsidence, it is likely that the summits of kuhn ø, where palaeogene basalts are present, were also buried below this palaeogene cover. the cover was, however, possibly thinner than those on clavering ø. the absence of evidence for elevated end-eocene palaeotemperatures on kuhn ø may be due to the difficulty in resolving multiple cooling episodes from the afta data, particularly where these are closely spaced in both time and palaeotemperature. two events dominate the late-stage cooling of the afta data from kuhn ø: the mid-cretaceous c4 and late miocene c8 episodes for which palaeotemperatures in the summits are about 60–70 and 30–40°c, respectively. intermediate end-eocene palaeotemperatures between these ranges are not likely to be resolved by the afta data. on store koldewey, the end-eocene episode was also not resolved in the afta data. here, common palaeotemperatures for the mid-cretaceous c4 episode are 85–95°c at sea level and those for the late miocene episode are 60–70°c. again, this implies that end-eocene palaeotemperatures within or below this range are not likely to be resolved by the afta data. removal of a cover of for instance 1 km of palaeogene basalts and eocene sediments in the end-eocene episode would not be detected. the end-eocene event was not resolved in the afta data for milne land. here, and on the adjacent island of gåseland, the ups defines a clear erosional truncation of both palaeogene basalts and precambrian basement (fig. 7). the former presence of a thicker cover of basalts on milne land is supported by a study of zeolite zones in the basalts there (larsen et al. 1989). these authors suggested that the absence of the shallow and less altered zeolite zones on milne land likely reflected removal of these zones by regional erosion after tilting of the basalts. larsen et al. (1989) assumed that the palaeogeothermal gradient was 40°c/km and concluded that about 900 m of basalts were removed across milne land. in summary, afta data show that a cover of up to 2 km of palaeogene basalts and eocene sediments were deposited south of scoresby sund and in the clavering ø region. as the cover was deposited during post-rift thermal subsidence, it must have been continuous along the entire margin, even reaching the inland nunataks as well as milne land. north of wollaston forland, the removal of this palaeogene cover cannot be resolved in the afta kap dalton milne land jameson land liverpool land kap brewster b l o s s e v i l l e k y s t 50 km 70°n 25°w scoresby sund end-eocene c6 palaeotemperatures kap brewster, gc1019-32 bopladsen fm, early eocene (post-basalt): >90°c >2.8 km eocene cover kap dalton, gc1016-21, 22 igertivâ fm, early eocene (intra-basalt): >75–100°c, 60–100°c 2.2–3.2 km eocene cover pre-devonian rocks afta samples carboniferous–cretaceous sediments vindtop fm (middle miocene) palaeogene basalts quaternary sediments cenozoic sediments fig. 27 post-basalt heating of lower eocene sediments at blosseville kyst (larsen et al. 2013; japsen et al. 2014). end-eocene c6 palaeotemperatures in three samples (yellow circles) correspond to burial below a c. 2.5 km thick eocene succession assuming a palaeothermal gradient of 25°c/km and a surface temperature of 20°c. the deposition of this succession took place over an interval of c. 12 myr (between 48 and 36 ma). http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 51 of 154 www.geusbul let in.org data, probably due to the high palaeotemperatures during the mid-cretaceous and late miocene episodes. 6.1.3 end-eocene differential, vertical movements end-eocene palaeotemperature profiles from vertical transects in appendix 2.3 show differences between various locations, suggestive of differential vertical movements (i.e. amounts of exhumation). figure 28 illustrates thermal-history constraints estimated at summit level on hudson land derived from palaeothermal constraints from afta in two vertical transects, vt3 and vt4, integrated with the geological record (locations on fig. 11b). samples from vt3 are from outcrops of carboniferous sandstones, whereas samples from vt4 are from outcrops of crystalline basement. vt3 is located c. 20 km north-west of vt4. both transects are east of the pdmf (fig. 3). the storelv fault, south-east of vt4, separates down-faulted, palaeogene basalts from basement (fig. 29). no palaeogene basalts are present above these transects. however, since basalts are preserved on nearby summits, it is likely that summit rocks on hudson land were near the surface during extrusion of the basalts in the early eocene. palaeotemperatures extrapolated to present-day summit elevations are higher for vt4 than for vt3 during the earliest cretaceous c3 and end-eocene c6 episodes, but not during the late miocene c8 episode. these differences are also seen when comparing the end-eocene pse over the short distance between vt3 and vt4 (fig. 26). there are two possible explanations: differential vertical movements of several hundreds of metres between the two transects after the onset of c6 cooling or higher heat flow prior to episode c8 at vt4 – or a combination of both. comparison of the profiles in appendix 2.5 (combined with the proximity of the two locations) appears to rule out a significant difference in heat flow between the two locations. fieldwork in 2010 confirmed the presence of the storelv fault in the area, so we interpret the different heating below upper carboniferous to middle triassic sediments ?? ? ? palaeogene basalts in the summits of clavering ø di�erence in palaeotemperatures between vt3 and vt4 levelled out after c6 block rotation and faulting ? upper carboniferous sandstones fig. 28 thermal-history for the summits of hudson land, based on palaeothermal constraints from afta, the geological record and stratigraphic landscape analysis (see discussion in appendix 3). palaeotemperatures are estimated for individual samples and projected to the elevation of the local summits. no palaeogene basalts are present in the summits above vt3 and vt4, but basalts are present at summits on clavering ø and hold with hope. the palaeotemperatures estimated are higher for vt4 than vt3 during the earliest cretaceous c3 and end-eocene c6 episodes, but not during the late miocene c8 episode. this difference is primarily due to differential exhumation between the two transects during the c6 episode caused by end-eocene block rotation and faulting. post-basalt faulting also occurred just east of vt4 along the storelv fault (fig. 29). shaded boxes: vt3. dashed boxes: vt4. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 52 of 154 www.geusbul let in.org thermal histories at vt3 and vt4 as caused by block rotation during the end-eocene c6 episode of the eastern part of hudson land (figs 28, 29; guarnieri 2015). the easternmost part of the block (vt4) thus underwent more uplift and erosion than farther inland (vt3) where carboniferous sediments are preserved. during this rotation, the basalts were downfaulted to the east of the storelv fault. finally, by the late miocene, the two blocks were locked in their present position as demonstrated by the similar palaeotemperatures for the two transects at that time. the storelv fault is located just west of loch fyne that hosts the major, post-basalt fault zone in the hold with hope and gauss halvø region, and this fault zone has an eastward throw of over 500 m (see location in figs 2, 29; upton et al. 1980). we infer that the faulting along loch fyne also occurred in the late eocene. the magnitude of end-eocene block rotation on hold with hope is 600 m as estimated by the difference between end-eocene pse at vt4 and vt3 (fig. 26), which is in good agreement with the estimate of upton et al. (1980). the end-eocene pse also varies over a short distance along the north coast of ymer ø, from 2.9 km to c. 3.5 km a.s.l. from west to east (fig. 26). however, the late miocene pse also varies over a short distance there, so it is possible that a part of the offset across the pdmf occurred in the late miocene (section 6.2.2). where afta data reveal cooling during the c6 episode they correspond closely to the rift basins east of the pdmf, from jameson land to wollaston forland. we interpret this to imply that exhumation had a greater magnitude across the rift basins, and that they consequently suffered more inversion in this episode than regions west of the pdmf where the effects of the mid-cretaceous c4 episode dominate. consequently, we infer that the pdmf was active during the end-eocene uplift phase. this interpretation is supported by bütler’s (1955) observation that tertiary dykes on ymer ø and on geographical society ø are offset along the pdmf; see also hartz et al. (2006). mid-cenozoic compressional deformation affected the east greenland shelf (tsikalas et al. 2005), providing further support for the scenario described here. 6.2 miocene burial and exhumation 6.2.1 development of the present-day topography within the last 10 myr the late miocene c8 pse is between 2 and 2.5 km a.s.l. across most of the study area, with values reaching 3 km or more east of the pdmf on geographical society ø, traill ø, jameson land and liverpool land (fig. 30). the thickness of the cover at that time is calculated from the late miocene palaeotemperature, assuming a surface temperature of 15°c and a palaeogeothermal gradient of 25°c/km (30°c/km for eastern traill ø, zones 2 and 3, and vt14). the late miocene pse is calculated for samples where the mid-cretaceous c4 or the end-eocene c6 episodes were resolved, suggesting that c8 palaeotemperatures are defined with confidence. only one sample in the interior highlands stands out as anomalous relative to the general picture (gc1077-62). this is probably a result of unresolved events in the afta data, or else represents a statistical outlier. in the interior highlands, the late miocene pse corresponds to the general present-day summit level in these mountains and is close to the elevation of the ups loch fyne nordhoek bjerg storelv fault basement palaeogene basalts fig. 29 contact between basement and palaeogene basalts along the storelv fault, hudson land (looking west). samples along vertical transect vt4 were taken in basement rocks just north of this locality. afta data from vt4 and vt3 in this region indicate faulting during the end-eocene c6 episode between these locations (fig. 28). photo: p. guarnieri (see also guarnieri 2015). http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 53 of 154 www.geusbul let in.org (figs 9, 22, 30). following the definition of the pse, this implies that the ups and the present-day summit level correspond to the late miocene surface. therefore, late miocene cooling in samples at elevations below the summits (i.e. below the ups) in this region can be explained in terms of removal of the rock below the summit (ups) by incision of the present-day relief. since incision is the result of uplift of the present-day mountains, formation of the relief since 10 ma implies, in turn, that uplift began at c. 10 ma. at coastal locations, stretching from store koldewey to liverpool land, the late miocene pse is higher than the summits that define the lps, and also higher than the ups where its presence can be inferred. further, late miocene palaeotemperature profiles in coastal regions show increasing offsets to higher temperatures compared to inland locations (fig. 31). in coastal locations, the summits must have been buried by a considerable cover when cooling began at c. 10 ma. the results from clavering ø (vt1; appendix 2.1) are a good example of this interpretation. palaeogene basalts are present in the summits that coincide with the lps at pdmf reactivated in cenozoic time 1 23 4 4a 5 6 8 9 10 14 11 12 13 15 17 16 3.5–3.9 3.0–3.4 2.5–2.9 2.0–2.4 1.9–2.2 622.0 1.8–2.3 3.2 3.0 <3.1 late miocene projected surface elevation (km a.s.l.) 2.2 2.2–2.5 2.4 2.0–2.6 2.32.3 3.3 3.3–3.5 2.2 2.0–2.3 2.9–3.2 2.3–3.5 unres. crt. unres. crt.: unresolved cretaceous event unres. crt. unres. crt. 2.3–2.5 unres. crt. unres. crt. 3.2–4.2 unres. crt. 2.1–2.9 2.6 2.6 traill ø clavering ø wollaston forland ymer ø geographical society ø hope with hope milne land store koldewey jameson land liverpool land kong oscar fjord vertical transect sample 1 50 km fig 6x7 76 °n 74 °n 72 °n 70 °n 70 °n 72 °n 74 °n 76 °n 30°w 20°w 30°w 20°w scoresby sund fig. 30 common projected surface elevations (pse; fig. 25) above present sea level for the late miocene episode for groups of samples outlined by thin black lines. pse is calculated for a surface temperature of 15°c and a palaeogeothermal gradient of 25°c/km (30°c/ km for eastern traill ø, zones 2 and 3 (fig. 32), and vt14). for a clearer overview, single outliers within 300 m of the common pse for a group of samples, are not shown. pses for samples in which the mid-cretaceous c4 episode is not resolved are not shown, but their locations are indicated (not for vertical transects). due to the uncertainty in the sample elevation, data from the gc103, gc159 and gc202 reports were not used. data for gc1077-77 were of poor quality and not used. two main pse levels are shown: c. 2.0 to 2.5 km a.s.l. in the north and western interior; 3 km a.s.l. or more in the south-east. pse values in stauning alper are around 2.6 km. the jump of pse values across the post-devonian main fault (pdmf), is indicative of miocene reactivation of the fault (e.g. along the north coast of ymer ø). only one sample in the interior highlands stands out as a significant outlier relative to the general picture (gc1077-62). http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 54 of 154 www.geusbul let in.org an elevation of c. 1.2 km a.s.l. (fig. 9). taking constraints from all vt1 samples into account, the late miocene pse becomes 2.4 km a.s.l. implying that magnitude of rock uplift since 10 ma (fig. 30). the cover above the basalts on clavering ø in the late miocene are therefore estimated to be c. 1.2 km, consisting of palaeogene basalts and younger sediments. there is no stratigraphic evidence that could refine the age of these cenozoic rocks, but the elevation of the ups – a measure of rock uplift since the formation of the surface – provides some constraints. the ups is not preserved on clavering ø, but it is present to the south-west on western hudson land, where its elevation is 1.8 km a.s.l. (fig. 9). this elevation is 700 m above that of the lps on clavering ø, which is a typical distance between the ups and lps across most of the region. it therefore seems reasonable to extrapolate the ups elevation to clavering ø. the late miocene pse on clavering ø (2.4 km a.s.l.) is 600 m above the elevation of the ups (1.8 km a.s.l.), suggesting that the ups was buried below 600 m of rock when uplift began at 10 ma. the ups was eroded to sea level as a result of uplift and erosion beginning 35 ma. so, it seems reasonable to conclude that the additional 600 m of rock that was present above the ups was mainly miocene sediments, deposited between the creation of the surface (some time prior to 10 ma) and the onset of late miocene uplift (beginning 10 ma). applying a higher palaeogeothermal gradient would reduce the estimate of the late miocene pse and of the thickness of the removed cover, but some thickness of cover would still be required (see further discussion in appendix 2). we therefore suggest that there is good evidence that up to 600 m of mainly miocene sediments accumulated in the coastal zone around clavering ø. to generate the accommodation space necessary for the accumulation of the miocene sediments in the coastal zone, this area must have subsided, in contrast to the areas further inland. this must have happened during a period of crustal extension prior to the onset of late miocene uplift. the high values of the late miocene pse along the coast in turn indicate that rock uplift since 10 ma was higher than further inland. this implies that the coastal zone must have been in compression during the late miocene, which agrees with observations both onshore and offshore north-east greenland (see section 2.4.1; price et al. 1997; hamann et al. 2005; døssing et al. 2016). the relatively smooth variation of the late miocene pse across wide areas (fig. 30), demonstrates that the topography in the late miocene was characterised by a low-relief surface. in the interior topography was dominated by the ups (which had been graded to base level following post-basalt uplift and erosion) and in the el ev at io n (k m ) 0.5 25°c/km 1.0 0.0 1.5 palaeotemperature (°c) a 20 40 60 80 100 120 140 160 1800 2.0 vt1: clavering ø (w) vt2: hold with hope (nw) vt3: hudson land (n) vt4,4a: hudson land (s) vt5: geographical society ø vt6: traill ø (s) vt8: milne land vt9: kuhn ø vt10: payer land vt11: stauning alper vt12: liverpool land vt13: jameson land vt14: giesecke bjerge vt15: traill ø (e) vt16: dombjerg vt17: clavering ø (ne) 30°w 20°w 30°w 20°w 76 °n 74 °n 72 °n 72 °n 70 °n magnitude of late miocene cooling from vt1–vt17 highest traill ø high medium lowest b 99 1 23 4 14 10 13 8 11 16 17 5 15 6 7 12 fig. 31 constraints on late miocene c8 palaeotemperatures from vertical transects. a: vertical transects for four groups of data (location in fig. 31b) define sub-parallel trends offset to higher temperatures. the largest offsets may define a higher palaeogeothermal gradient (vt6 and vt15 from traill ø). b: relative variation in exhumation across the study region. data from vt7 are not discussed due to uncertainties in elevation. vt: vertical transect. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 55 of 154 www.geusbul let in.org coastal zone by the top of lower to middle miocene sediments that accumulated during a phase of subsidence prior to late miocene uplift. in any case, the late miocene land surface cannot have been much above sea level. consequently, the altitude of the late miocene pse is a measure of post-late miocene uplift of these surfaces, prior to their dissection. 6.2.2 late miocene differential, vertical movements in the southern part of the study area (south of c. 73°n) there is good east–west data coverage. here, the late miocene pse increases from about 2 km a.s.l. across wide areas of the interior highlands to about 3 km a.s.l. in the coastal zone (fig. 30). part of this increase could be due to a higher palaeogeothermal gradient in the coastal zone. at some locations this change takes place over a short distance. specifically: • along kong oscar fjord. values increase from c. 2.2 km a.s.l. in the inner part of the fjord to 3.5 km a.s.l. along the northern coast and to 3.2 km a.s.l. on the southern coast. the palaeogeothermal gradient there is constrained by the afta data to increase from about 25 to 30°c/km along the fjord (fig. 32), so the observed difference in pse cannot be explained solely by different heat flow and differential exhumation is clearly a dominant factor. only by assuming a palaeogeothermal gradient of 50–60°c/km can the late miocene pse for vt6 be reduced to about 2 km a.s.l., but this is not allowed by the data. • along the north coast of ymer ø. a similar increase is observed, from c. 2.3 km a.s.l. in the west to 3.3 km a.s.l. in the east. analogous with those along kong oscar fjord, we interpret these differences as largely due to differential uplift and erosion. • jameson land and liverpool land. whereas the values over most of the southern interior are around 2 km a.s.l., the pse on jameson land and liverpool land range from 2.9 to 3.2 km a.s.l., respectively. thus, some coastal areas were subjected to deeper exhumation after 10 ma than areas farther inland. the strong lateral difference in the estimated rock uplift over short distances across the pdmf, indicates that late miocene, differential movements possibly occurred across this fault. in section 6.2.1, we showed that a thick cover of miocene sediments accumulated in the coastal zone during a period of crustal extension prior to late miocene uplift during a compressional phase. it is thus possible that both burial and subsequent exhumation in the coastal zone were guided by differential movements along the pdmf. 6.3 cenozoic events in the traill ø area late miocene vts that plot furthest to the right in fig. 31a are from traill ø, in proximity to late eocene intrusive bodies where end-eocene palaeotemperatures define an elevated palaeogeothermal gradient (section 5.8). palaeotemperatures in end-eocene c6 and late miocene c8 episodes define local maxima around the palaeogene igneous centres on traill ø (fig. 14; price & witham 1997; price et al. 1997). in addition, this area is one of the few parts of the study area where early pliocene cooling episode c9 is recorded. to further investigate vertical and horizontal variation of palaeotemperatures, we divided the region around traill ø into four zones and plotted palaeotemperatures in episodes c6, c8 and c9 against elevation in each zone (fig. 32). zone 3 represents the igneous centre and zone 0 represents background effects at some distance from the igneous centre. in both zones 2 and 3, palaeotemperature constraints for the c6, c8 and c9 episodes show a clear separation into 3 separate linear profiles, where palaeotemperatures increase with decreasing elevation (representing palaeoburial) in each episode. higher palaeotemperatures in each episode in these zones allow resolution of the separate late miocene and early pliocene cooling episodes. in contrast, constraints in samples from zones 0 and 1 do not allow clear discrimination of three separate events. here, some samples provide palaeotemperature constraints, which would be more consistent with a profile offset towards greater depths compared to the majority of samples, and are therefore attributed to the lower magnitude, early pliocene c9 episode. in zones 1, 2 and 3, c6 palaeotemperatures increase towards the igneous centres at the eastern extremity of traill ø (zone 3). although not definitive, the profiles for episode c6 in fig. 32 are consistent with an eastwards-increasing palaeogeothermal gradient from c. 25°c/km in zone 1 to c. 45°c/km (or possibly higher) in zone 3, as might be expected due to the presence of the igneous centres. late miocene c8 palaeotemperatures also show a clear increase from west to east from zones 0 to 3. the profiles (drawn by eye) through the constraints suggest an increasing palaeogeothermal gradient from 25°c/km in zones 0 and 1 to c. 30°c/km in zones 2 and 3. this indicates that some residual elevated heat flow related to the intrusive activity may have remained c. 25 myr after intrusion. the progressive offset in the late miocene profiles towards higher palaeotemperatures from zones 0 to 3 reflects progressively deeper levels of palaeoburial prior to the onset of cooling, which we interpret as due predominantly to exhumation. early pliocene c9 palaeotemperatures in zones 2 and 3 are consistent with linear profiles defined by a http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 56 of 154 www.geusbul let in.org 16 17 18 14 9 10 4 1 11 2 6 29-33 35-37 42 5051 52,54,55 5 6 yø a b c d e gsø tø 15 zone 3 zone 2 zone 1 zone 0 15 19 16,17 18 el ev at io n (k m ) 0.5 1.0 0.0 1.5 palaeotemperature (°c) 20 40 60 80 100 120 140 160 1800 c. 5 ma: 25°c/km c9 c. 10 ma: 30°c/km c8 37–35 ma: 45°c/km c6 zone 3 gc1077-52 gc1077-55 gc1104-18 gc1077-54 gc522-1 gc522-11 gc522-4 22–7 ma17–0 ma 20–0 ma 25–0 ma 14–4 ma 28–10 ma 22–5 ma 5–0 ma 13–0 ma 8–0 ma el ev at io n (k m ) 0.5 1.0 0.0 1.5 palaeotemperature (°c) 20 40 60 80 100 120 140 160 1800 c. 5 ma: 25°c/km c9 c. 10 ma: 30°/km c8 37–35 ma: 40°c/km c6 zone 2 gc1104-19 gc1104-17, 16,15 gc1016-32 gc1016-29 gc1016-33 gc522-10 gc1016-3010–0 ma 12–0 ma 30–0 ma 30–0 ma 7–0 ma 20–10 ma 12–3 ma 43–10 ma 13–2 ma 12–4 ma 27–5 ma el ev at io n (k m ) 0.5 1.0 0.0 1.5 palaeotemperature (°c) 20 40 60 80 100 120 140 160 1800 c. 5 ma: 25°c/km c9 c. 10 ma: 25°c/km c8 37–35 ma: 25°c/km c6 zone 1 35–0 ma 32–2 ma 20–0 ma 15–5 ma 15–5 ma 10–3 ma gc1016-37 gc1016-36 gc1077-50 gc1077-51 gc522-9 gc1016-35 el ev at io n (k m ) 0.5 1.0 0.0 1.5 palaeotemperature (°c) 20 40 60 80 100 120 140 160 1800 16–0 ma 30–0 ma40–0 ma 12–0 ma 15–0 ma 15–5 ma gc1016-14, 16,17,18 gc522-6 gc522-2 zone 0 c. 5 ma: 25°c/km c9 c. 10 ma: 25°c/km c8 fig. 32 palaeotemperature constraints from afta in samples from four zones of the traill ø region. a: sample locations and zonation (zones 0–3). b–e: palaeotemperature profiles vs. elevation a.s.l. constraining three palaeothermal episodes (c6, c8 and c9). palaeotemperature constraints for c5 in zone 2 and c6 in zone 0 are not shown. constraints in each zone are separated into a series of linear profiles where palaeotemperatures increase with decreasing elevation (representing palaeoburial). this allows improved attribution of constraints to the discrete regional cooling episodes, and this information was employed to attribute constraints to individual episodes (appendix 1.7, also available online in supplementary file s4). c6 palaeotemperatures and palaeogeothermal gradients increase towards the igneous centres at the eastern extremity of traill ø (zone 3). c8 palaeotemperatures also increase towards eastern traill ø, but the palaeogeothermal gradient only shows a minor variation in that direction, indicating depth of burial was the dominant factor for increasing palaeotemperatures. gsø: geographical society ø. tø: traill ø. yø: ymer ø. circles and quadrants with numbers: afta sample(s) and vertical transects, respectively (see fig. 11). http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 57 of 154 www.geusbul let in.org palaeogeothermal gradient of 25°c/km. in zone 3, afta data in the three ‘deepest’ samples (i.e. those closest to sea level) clearly resolve both the c8 and the c9 cooling episodes, due mainly to higher c8 palaeotemperatures of around 100°c or more, where the technique is most sensitive. similar comments apply to the ‘deepest’ sample in zone 2. in samples from higher elevation, the afta data do not allow resolution of two separate episodes. however, comparing results from the two ‘highest’ samples in zone 2, define the most recent cooling episode from 80–90°c beginning between 20 and 10 ma (sample gc1016-30), and cooling from 45–75°c, beginning between 10 and 0 ma (sample gc1104-19). these two events plots on the profiles characterising the c8 and c9 episodes, respectively. these results show that two discrete episodes of cooling have occurred in the region around traill ø within the last 10 myr. the failure to resolve discrete c8 and c9 episodes in areas other than traill ø is probably due to a lower magnitude of pliocene exhumation across the wider region, such that palaeotemperatures were too low to be defined from afta (i.e. <50°c). however, at the eastern extremity of traill ø, significant pliocene exhumation combined with high c6 and c8 palaeotemperatures allows resolution of the separate the c8 and c9 cooling episodes. 6.4 stauning alper at stauning alper, elevations reach 2.8 km a.s.l. there is significant alpine relief, and it is one of the most dissected areas in north-east greenland. according to the stratigraphic landscape analysis (chapter 3), the ups defines the summits of stauning alper. the elevation of the ups around stauning alper forms a domal structure with maximum ups elevations of c. 2.3 km a.s.l. (fig. 9). this relief is probably not related to differences in rock type because similar bedrock (mainly crystalline basement) occurs in nearby areas at almost the same elevation (e.g. renland and milne land). however, the ups is very well preserved in these areas. stauning alper clearly experienced more uplift than the surrounding areas after the formation of the ups, and these mountains thus have a tectonic history that deviates from the regional pattern. for example, uplift here may have started earlier than in the rest of the region, leaving more time for the destruction of the ups, and thus stronger valley incision and relief rejuvenation. further erosion may have been by cirque glaciers that formed at high elevation. the early miocene c7 episode is identified across northern jameson land, just east of stauning alper. the onset of uplift in stauning alper may have begun in the early miocene rather than in the late miocene. an uplift component due to the c7 episode does seem to be a reasonable interpretation, given the emplacement of the massive igneous bodies during the latest oligocene at, for example, malmbjerget, at the eastern edge of stauning alper. more work, involving detailed sampling of known igneous bodies and their elevation, would be required to resolve these remaining uncertainties. 6.5 a stepped topography formed after three phases of uplift and incision the palaeothermal constraints from afta, combined with the denudation chronology defined from the stratigraphic landscape analysis, as well as the geological record, provide clear constraints on the history of vertical motions of the east greenland margin. based on stratigraphic landscape analysis, bonow & japsen (2021, this volume) identified two peneplains across northeast greenland and established a relative denudation chronology for their formation after break-up and onset of drift (fig. 33; section 3.2): 1. the ups (typically at 2 km a.s.l.) cuts across palaeogene basalts and older rocks (fig. 33b). a first phase of uplift in post-basalt times led to the formation of the ups by erosion to a base level defined by the atlantic ocean. the ups had formed before the extrusion of the vindtop formation in the mid-miocene (bonow et al. 2014). 2. the lps (typically at 1 km a.s.l.) is defined by broad valleys below the ups in the milne land region and by broader zones along the coast (fig. 33d). a second phase of uplift led to the formation of the lps by river incision to base level below the uplifted ups. 3. a third phase of uplift raised the ups and the lps to their present elevations (fig. 33f). the presence of two elevated peneplains across the region thus requires three phases of post-breakup uplift. our synthesis of afta data identifies three regional phases of uplift and erosion after break-up (section 4.5), during (1) the end-eocene c6 episode (fig. 33a); (2) the late miocene c8 episode (fig. 33c) and (3) the early pliocene c9 episode (fig. 33e). we suggest that each of these episodes of uplift and erosion led to the formation of distinct steps in the landscape (fig. 34) as follows: 1. following end-eocene maximum burial at c. 35 ma, uplift and erosion led to formation of the ups as a peneplain graded to sea level. 2. late miocene uplift at c. 10 ma led to formation of the lps by incision and valley widening below the uplifted ups. in parts of the coastal zone, a phase of miocene burial preceded the uplift (section 6.2.1). 3. early pliocene uplift at c. 5 ma led to incision of the http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 58 of 154 www.geusbul let in.org 2100 1900 1700 1500 1300 1100 900 700 500 300 100 0 elevation (m a.s.l.) a. end-eocene c6 uplift c6 cooling unresolved by afta data? b. upper planation surface (ups) uplift denudation c. late miocene c8 uplift from afta data from afta data from afta data d. lower planation surface (lps) e. early pliocene c9 uplift f. present relief c8 c6 ups u ps d es tr oy ed ? u ps d es tr oy ed ? lps c9 cooling unresolved by afta data? c9 30°w 20°w 30°w 20°w 76 °n 20°w30°w 20°w30°w 76 °n 74 °n 72 °n 70 °n 20°w30°w 20°w30°w 76 °n 74 °n 72 °n 70 °n 20°w30°w 20°w30°w 76 °n 74 °n 72 °n 70 °n 20°w30°w 74 °n 72 °n 70 °n 76 °n 20°w30°w 74 °n 72 °n 70 °n fig. 33 cenozoic uplift events and relief formed by denudation in the study area. a: area where uplift during the end-eocene c6 episode is recognised (based on fig. 14). b: mapped extent of the upper planation surface (ups; based on fig. 9). c: area where uplift during the late miocene c8 episode is recognised (based on fig. 14). d: mapped extent of the lower planation surface (lps; based on fig. 9). e: area where uplift in the early pliocene c9 episode is recognised (based on fig. 14). f: present-day relief (amante & eakins 2009). http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 59 of 154 www.geusbul let in.org present-day valley floors below the uplifted lps and to formation of the present-day landscape. therefore, this is the third of the cenozoic episodes that correlates with the regional post-breakup unconformity as discussed in sections 5.7 and 5.9. with this sequence of events, it is possible to explain the presence of two elevated peneplains that formed in post-basalt times with the three regional events of uplift and exhumation since break-up defined by afta. however, the areal extent of the two peneplains and of the regions where the three uplift events have been recognised, only partially overlap (fig. 33). the ups primarily extends across basement terrains in the interior highlands, typically at elevations about 2 km a.s.l. (fig. 33b). north of 74°n, where the greenland ice sheet and extensive glaciers are closer to the coast than to the south, the ups is only defined in narrow patches along the ice cover or as nunataks. the lps defines the summits around 1 km a.s.l. in a c. 100 km wide band along the coast, but it continues as elevated palaeovalleys below the ups in the interior highlands (fig. 33d). the two denudation surfaces and their remnants are thus both identified across most of the interior highlands south of 74°n as well as across hudson land and western traill ø. farther north, there is only limited overlap between the areas where the surfaces are identified. of the three palaeothermal episodes in post-breakup time, only the late miocene episode is identified across the entire region. there is, however, evidence to suggest that the other two episodes of uplift and erosion also acted on a regional scale. the end-eocene episode is only recognised in the eastern part of the study area (fig. 33a), but this area overlaps with the extent of both the ups and the lps in the inner part of the coastal zone from traill ø to hudson land (figs 33b, d). furthermore, we argued in section 6.1.2, that an eocene cover is likely to have been present across wide parts of the study area and then removed without necessarily being detected in the afta data. we therefore find that the continuity of the ups from this central region into the interior supports the conclusion that the ups there was also formed after end-eocene uplift and exhumation. furthermore, the continuity of the lps from this central region towards the north suggests that these areas were also affected by end-eocene uplift and formation of the ups, followed by a phase of uplift and formation of the lps. apparently, in these northerly regions, eocene burial and exhumation were relatively minor compared to the magnitudes of the mid-cretaceous and late miocene events. other evidence for regional uplift and exhumation includes the regional nature of the late miocene event comparable to that of the lps (figs 33c, d), and that present-day elevation of the lps is due to early pliocene c9 uplift. whereas the lps is recognised at an elevation of around 1 km a.s.l. across most of the region, the early pliocene episode was only identified in afta data from a limited area (fig. 33e). however, as argued previously, we suggest that the pliocene event also affected the wider region. interestingly, the continuity of the lps on both sides of the pdmf in the south-western part of the study area (section 3.1), indicates that this fault system was 35 ma 10 ma c. 5 ma post-basalt sediments palaeogene basaltsmiocene basalt pre-basalt sediments basement lpslps fig. 34 block diagrams illustrating the post-rift development of the east greenland margin since cenozoic maximum burial below a palaeogene succession of volcanics and sediments by end of the eocene. three events of uplift and exhumation (end-eocene c6, late miocene c8 and early pliocene c9) led to the formation of the present-day landscape with two elevated plains, the upper and lower planation surface (ups and lps, respectively). the presence of the mid-miocene basalts flows of the vindtop formation (storey et al. 2004; figs 21, 27) on the ups shows that this surface was formed in pre-glacial times (bonow et al. 2014). http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 60 of 154 www.geusbul let in.org not active during the early pliocene episode. this is in contrast to the movements of the pdmf that we have inferred for the end-eocene and late miocene episodes further north (sections 6.1.3 and 6.2.2, respectively). we conclude that the presence of the ups and lps as two major steps in the landscape of north-east greenland can be explained as a consequence of sequential uplift and incision beginning in the end-eocene, late miocene and early pliocene. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 61 of 154 www.geusbul let in.org in this chapter, we present reconstructed thermal histories for locations that reveal the regional variation in palaeozoic and mesozoic tectonic processes. whereas deeper burial and exhumation control the regional variation in thermal histories, for the most part, evidence for vertical offsets across faults is also present. we also discuss the age and thickness of the covers that were present prior to the pre-cenozoic episodes of uplift and erosion, based on information presented in chapters 4 and 5. 7.1 thermal-history reconstructions, store koldewey thermal-history solutions derived from afta in three samples of caledonian basement from store koldewey are shown in fig. 35. many basement samples from the region north of 75°n share essentially similar thermal histories, undergoing exhumation and cooling below c. 100°c in the late carboniferous c0 episode, sometime between 320 and 300 ma. furthermore, deep basement levels were exhumed in the middle devonian prior to deposition of old red sandstones (gilotti & mcclelland 2008). 7. palaeozoic–mesozoic thermal histories deposition of sediments deposition of sediments deposition of sediments ? ? ? ? 20 40 60 80 100 120 140 500 te m pe ra tu re (° c ) 0100200300400 0 20 40 60 80 100 120 140 500 te m pe ra tu re (° c ) 0100200300400 0 20 sil. dev. carb. perm. trias. jura. cret. cen.ord. sil. dev. carb. perm. trias. jura. cret. cen.ord. sil. dev. carb. perm. trias. jura. cret. cen.ord. 40 60 80 100 120 140 500 te m pe ra tu re (° c ) time (ma) gc1077-16 gc1077-83 5 km store koldewey gc1077-102 0100200300400 0 devonian unroo�ng devonian unroo�ng ? ? ? ? devonian unroo�ng ? c8 c6c4c3c2c1c0 c8 c6c4c3c2c1c0 c8 c6c4c3c2c1c0 76 °1 5' n 76 °1 5' n 19°w 19°w fig. 35 variation in thermal histories in caledonian basement samples from store koldewey. coloured vertical bars define the onset of regional cooling episodes (table 1). coloured boxes define the palaeotemperature–time constraints for individual samples. the 95% uncertainty limits on the onset of cooling are quite wide for some samples. attribution to regional episodes is based on a comparison of samples within a restricted region. map location in fig. 11. ord.: ordovician. sil.: silurian. dev.: devonian. carb.: carboniferous. perm.: permian. trias.: triassic. jura.: jurassic. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 62 of 154 www.geusbul let in.org 7.1.1 samples gc1017-16, -102, crystalline basement, near sea level basement samples gc1077-16 and -102, both from low elevations, cooled below c. 100°c in the late carboniferous c0 episode and were exhumed to the surface by the late jurassic (gc1077-16) and early cretaceous (gc1077-102), when sediments were deposited on top of the sampled basement. both sediments and basement were then reburied to reach maximum palaeotemperatures around 80 to 90°c, from which cooling/ exhumation commenced in the mid-cretaceous c4 episode (between 95 and 90 ma). vr and afta data from the overlying sedimentary successions corroborate evidence of this re-burial, as summarised in fig. 15. the palaeotemperature constraints for the mid-cretaceous c4 (c. turonian) episode for all samples from store koldewey are consistent with a pse (fig. 25) of about 2.8 km for a palaeogeothermal gradient of 25°c/km and a surface temperature of 20°c. for the samples of lower cretaceous sediments near sea level, this elevation corresponds to burial below a cretaceous cover of the same magnitude. the cover removed since the onset of the mid-cretaceous event must have been of aptian– turonian age as the youngest cretaceous sediments on store koldewey are of earliest aptian age (bjerager et al. 2020). intervening re-burial may have occurred following mid-cretaceous exhumation prior to the final phase of exhumation in the late miocene. 7.1.2 sample gc1017-83, crystalline basement, near sea level in contrast to samples gc1077-16 and -102, basement sample gc1077-83 did not cool below 100°c until the early jurassic (c. 180 ma), although the subsequent history is very similar. since sample gc1077-83 is located between samples gc1077-16 and -102, the early jurassic cooling though c. 100°c in sample gc1077-83, must reflect major faulting and tectonic offsets in this region. significant faulting must have taken place prior to the mid-cretaceous as nearby samples show very different thermal histories prior to that episode (e.g. samples gc1077-8 and gc1104-3). although no constraints on palaeogeothermal gradients are available from this region, results from locations to the south suggest that the thermal histories shown in fig. 35 reflect primarily the history of sequential exhumation/burial/exhumation. compare this with the thermal-history diagram for basement samples from store koldewey in appendix 3.12. 7.2 thermal-history reconstructions of pre-mesozoic terrains there is broad variation between thermal histories of palaeozoic sedimentary and intrusive rocks and crystaline basement across the region between 70 and 74°n (fig. 36). the basement was exhumed from great depth in the middle devonian prior to deposition of old red sandstones (gilotti & mcclelland 2008). here, we review a representative selection of samples to illustrate this regional variation. 7.2.1 sample gc1016-62, crystalline basement, hudson land (123 m a.s.l.) this sample cooled from >120°c in the earliest cretaceous c3 episode, followed by end-eocene c6 cooling from 110–120°c and finally late miocene c8 cooling from 40–80°c. compare this with the thermal-history diagram for vt4 in appendix 3.4. note the difference in c3 palaeotemperature compared to nearby sample gc1016-55 from clavering ø (next section). 7.2.2 sample gc1016-55, carboniferous sediments, clavering ø (52 m a.s.l.) this sample cooled from palaeotemperatures >100°c in the middle triassic c1 episode (data from this sample do not definitively record this episode, which is inferred from adjacent samples). the sample subsequently cooled from 95–105°c in the earliest cretaceous c3 episode and from 65–75°c in the late miocene c8 episode. based on data from adjacent samples, it is likely that the region from which this sample was collected also underwent early jurassic c2 cooling (c2 in fig. 14). compare with the thermal-history diagram for vt1 in appendix 3.1. 7.2.3 sample gc1016-9, devonian sediments, ella ø (1 m a.s.l.) this sample cooled below 120°c in the early jurassic c2 episode and subsequently cooled from 60–75°c in the late miocene c8 episode. data from adjacent samples suggest that this sample most likely also underwent cooling in the earliest cretaceous c3 and the end-eocene c6 episodes, although these events cannot be resolved in data from this sample. 7.2.4 sample gc1016-45, ordovician intrusion, milne land (458 m a.s.l.) this sample cooled below 100°c in the middle triassic c1 episode. based on regional geological considerations, the intrusion was unroofed in the middle devonian and subsequently covered by sedimentary rocks prior to c1 cooling. the sample subsequently cooled from 95–110°c in the early jurassic c2 episode, which is also interpreted as representing exhumation. finally, this sample cooled from a late miocene c8 palaeotemperature of 20–90°c. the presence of nearby palaeogene basalts shows that the sample was close to the surface in the early eocene. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 63 of 154 www.geusbul let in.org 7.3 thermal-history reconstructions of mesozoic terrains the broad range of thermal histories for samples from mesozoic sedimentary rocks across the region is shown in fig. 37. representative examples are reviewed below. 7.3.1 sample gc1016-69, triassic sediments, hold with hope (94 m a.s.l.) mid-cretaceous c4 cooling across hold with hope is supported by data from this sample, which cooled from 95–100°c sometime between 175 and 80 ma. the adjacent sample, gc1016-38, cooled from similar palaeotemperatures between 120 and 20 ma. sample 69 was subsequently affected by late miocene c8 cooling from 60–80°c. the presence of palaeogene basalts nearby again suggests that this sample was close to surface temperature during the early eocene. adjacent samples suggest that sample 69 also most likely underwent some degree of end-eocene c6 cooling. compare with the thermal-history diagram for vt2 in appendix 3.2. 7.3.2 sample gc522-7, jurassic sediments, wollaston forland (150 m a.s.l.) this sample reached a maximum palaeotemperature of 85–95°c prior to end-eocene c6 cooling. the sample subsequently cooled from 20–75°c in the late miocene c8 episode. the presence of extensive palaeogene basalts on wollaston forland suggests that the sample was close to surface temperatures during the early eocene, suggesting relatively rapid heating to the end-eocene palaeothermal peak. data from adjacent samples also suggest that this sample may have undergone earliest 30°w 20°w 20°w 76 °n 72 °n 70 °n hudson land clavering ø milne land ella ø 100 km gc1016-55 within ~1 km of surface ? devonian unroo�ng gc1016-45 devonian unroo�ng gc1016-9 deposition ? deposition of sediments c0 c0 c0 ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? ? earlier episodes de�ned in adjacent samples depositiongc1016-62 time (ma) c8 c6c4c3c2c1 c8 c6c4c3c2c1 c8 c6c4c3c2c1 c8 c6c4c3c2c1 20 40 60 80 100 120 140 500 te m pe ra tu re (° c ) 0100200300400 0 20 40 60 80 100 120 140 500 te m pe ra tu re (° c ) 0100200300400 0 20 40 60 80 100 120 140 500 te m pe ra tu re (° c ) 0100200300400 0 20 40 60 80 100 120 140 500 te m pe ra tu re (° c ) 0100200300400 0 time (ma) sil. dev. carb. perm. trias. jura. cret. cen.ord. sil. dev. carb. perm. trias. jura. cret. cen.ord. sil. dev. carb. perm. trias. jura. cret. cen.ord. sil. dev. carb. perm. trias. jura. cret. cen.ord. fig. 36 variation in thermal histories in samples of pre-mezozoic age. coloured vertical bars define the onset of regional cooling episodes (table 1). coloured boxes define the palaeotemperature– time constraints for individual samples. the 95% uncertainty limits on the onset of cooling are quite wide for some samples. attribution to regional episodes is based on a comparison of samples within a restricted region. ord.: ordovician. sil.: silurian. dev.: devonian. carb.: carboniferous. perm.: permian. trias.: triassic. jura.: jurassic. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 64 of 154 www.geusbul let in.org cretaceous c3 cooling, although palaeotemperatures at this time were lower than those reached in the end-eocene palaeothermal maximum. 7.3.3 sample gc522-10, jurassic sediments, traill ø (230 m a.s.l.) this sample cooled below 120°c in the c6 episode. it subsequently cooled from 100–110°c in the c8 episode and finally from 60–90°c within the last 5 myr (during the early pliocene c9 episode). evidence of the prelate eocene history is overprinted by the severity of end-eocene heating, but earlier events almost certainly affected this location. compare this with fig. 32 and with the thermal-history diagram for vt6 in appendix 3.6. 7.3.4 sample gc202-11, triassic sediments, south-eastern jameson land (150 m a.s.l.) this sample cooled below 120°c in the early eocene c5 episode (beginning c. 55 ma). this is interpreted as reflecting contact or hydrothermal heating associated with palaeogene intrusive activity (see section 5.6). it seems likely that burial may have been continuous during the paleocene and eocene. this sample subsequently cooled from 95–105°c in the end-eocene c6 episode, probably due to the onset of exhumation of the continental margin at this time. finally, this sample cooled from a late miocene c8 palaeotemperature of 40–65°c, interpreted as the final onset of exhumation of the continental margin. 100 km jameson land wollaston forland traill ø hold with hope 20°w 20°w30°w 70 °n 72 °n 74 °n 76 °n gc202-11 ? ? gc522-10 ? ? deposition deposition deposition gc1016-69 maximum post-deposition gc522-7 ? deposition near surface at ~55 ma? near surface at ~55 ma? c8c6c4c3c2 c5 c8 c6c4c3c2 c8c6c4c3c2 c8c6c4c3c2 c9 ? ? ? near surface at ~55 ma? 20 40 60 80 100 120 140 250 te m pe ra tu re (° c ) 050100150200 20 0 40 60 80 100 120 140 250 te m pe ra tu re (° c ) 050100150200 0 20 40 60 80 100 120 140 250 te m pe ra tu re (° c ) 050100150200 0 20 40 60 80 100 120 140 250 te m pe ra tu re (° c ) 050100150200 0 time (ma) triassic jurassic cretaceous cenozoic triassic jurassic cretaceous cenozoic triassic jurassic cretaceous cenozoic triassic jurassic cretaceous cenozoic fig. 37 variation in thermal histories in samples of mesozoic age. coloured vertical bars define the onset of regional cooling episodes (table 1). coloured boxes define the palaeotemperature– time constraints for individual samples. the 95% uncertainty limits on the onset of cooling are quite wide for some samples. attribution to regional episodes is based on a comparison of samples within a restricted region. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 65 of 154 www.geusbul let in.org 7.4 palaeozoic–mesozoic removed covers it is possible to refine the heating and cooling (burial and exhumation) history for several localities by combining the interpretation of afta data in (1) basement samples with geological constraints from sediments that rest on basement and (2) sediment samples with their stratigraphic age. based on a regional comparison of estimated palaeogeothermal gradients (section 4.6), we convert palaeotemperatures to thickness of covers based on a gradient of 25°c/km for all pre-cenozoic episodes; see further discussion in appendix 2. we assume a palaeosurface temperature of 20°c. sample locations mentioned in the following sections are shown in fig. 11. the maps of the appropriate palaeotemperatures are presented in fig. 14. 7.4.1 cover prior to exhumation in the middle triassic episode (c1) in the southern interior, west of milne land, one sample of the mid-permian røde ø conglomerate (stemmerik & piasecki 2004) was >110°c when c1 cooling began (gc1016-5). consequently, a cover of more than 3.5 km of mid-permian – middle triassic sediments must have been present above this sample at the onset of c1 cooling. in the summits of western clavering ø, a sample of upper carboniferous sandstone reached a c1 palaeotemperature between 105 and 115°c (gc1016-49; vt1; appendix 3.1). this must have been caused by burial below a thick cover (>3 km) of upper carboniferous to middle triassic sediments. on germania land, where upper carboniferous outliers are present, samples that cooled from c1 palaeotemperatures of 100°c or more reflect burial below a thick upper carboniferous – middle triassic cover (e.g. gc1077-11; appendix 3.13; piasecki et al. 1994). stemmerik et al. (1993) used thermal maturity data from upper permian deposits from clavering ø to northern jameson land, to estimate burial below an overburden of 1.5–3 km. the former presence of a kilometre-thick upper carboniferous (locally permian) – middle triassic cover is thus documented at very disparate locations on both sides of the pdmf system. it seems likely that this cover was present across much of the study area prior to c1 exhumation. 7.4.2 cover prior to exhumation in the early jurassic episode (c2) the constraints on the thermal history for two adjacent samples of palaeozoic granite and middle jurassic sandstones from milne land provide insights into the magnitude of exhumation in the early jurassic c2 episode (gc1016-45 and -44, respectively; vt8). whereas the sandstones were deposited on the eroded granite at an assumed temperature of c. 20°c, afta data show that the basement sample was at c. 100°c when c2 cooling began. for any reasonable palaeogeothermal gradient, this palaeotemperature implies burial below a kilometre-thick cover of pre-middle jurassic rocks prior to c2 exhumation. on store koldewey, middle jurassic sediments rest on basement and reached a c2 palaeotemperature >100°c, which again requires removal of a kilometre-thick pre-jurassic and lower jurassic rocks during the c2 episode (gc1077-83; appendix 3.12). a kilometre-thick cover was thus present above outcropping pre-jurassic rocks when cooling began during the early jurassic, both west and east of the pdmf system. 7.4.3 cover prior to exhumation in the earliest cretaceous episode (c3) afta data from clavering ø suggest a succession of 2 km or more of carboniferous to lowermost cretaceous sediments was present there at the palaeothermal peak prior to earliest cretaceous c3 cooling (compare this with the thermal-history diagram for vt1, appendix 3.1). on store koldewey, lower cretaceous sediments rest on basement that reached an earliest cretaceous c3 palaeotemperature of 80–90°c (gc1077-104; fig. 15). this implies removal of about 2.5 km of cretaceous rocks during the c3 episode. elsewhere, most of the evidence for c3 palaeotemperatures are from samples of basement terrains that do not offer chronostratigraphic constraints on the age of the rocks removed during this episode. additional insight into the nature of the c3 episode can be obtained from the afta data from the north-west coast of hold with hope (vt2; appendix 3.2). here, samples gc1016-66 to 69 of triassic sandstone units cooled from almost 100°c in the earliest cretaceous episode. as these units are overlain by middle jurassic deposits, the high c3 palaeotemperatures must have been produced by burial beneath a kilometre-thick cover of middle to upper jurassic sediments. 7.4.4 cover prior to exhumation in the midcretaceous episode (c4) at several localities within the study area, middle jurassic to lower cretaceous sediments were heated to mid-cretaceous c4 palaeotemperatures of 80 to 100°c. this is the case on store koldewey where c4 palaeotemperatures reached 80–90°c for most samples (including basement and sedimentary rocks; appendix 3.12). vr values between 0.43 and 0.58% (equivalent to maximum palaeotemperature of 72 to 96°c) obtained for lower cretaceous sediments there (fig. 15) are consistent with the c4 palaeotemperatures from afta. this shows that these lower cretaceous sediments cooled from their maximum post-depositional palaeotemperatures in http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 66 of 154 www.geusbul let in.org this episode. these palaeotemperatures correspond to burial below a c. 2.5 km-thick cretaceous succession. similarly, c4 palaeotemperatures correspond to values derived from vr data in the clavering ø region (as far south as hold with hope) at some distance from palaeogene intrusive centres, where palaeotemperatures derived from vr data reflect end-eocene c4 values (c4 and c6 in fig. 16). middle jurassic sediments on kuhn ø onlap basement close to the summits (c. 1.2 km a.s.l.) where the basement was heated to a c4 palaeotemperature around 70°c (vt9, appendix 3.7). this palaeotemperature corresponds to burial of the summits of kuhn ø below a 2 km thick middle jurassic to mid-cretaceous cover when c4 cooling began (representing maximum burial of the mesozoic succession). this succession thus accumulated in the interval between c2 and c4 exhumation and thus included sediments of toarcian to turonian age; not least a cretaceous post-rift succession that is no longer present on kuhn ø. on jameson land, cretaceous sediments are largely absent and the c4 episode could not be resolved by afta data. this is probably due to the dominance of the c3 episode in this region. figure 38 illustrates the wide extent of the corresponding kilometre-thick middle jurassic to cretaceous (aalenian–turonian) deposits present prior to c4 exhumation. the cover included the post-rift succession deposited in a broad thermal sag after rift-climax at the jurassic–cretaceous transition (surlyk 1978b, 1990). 20°w 20°w 30°w 30°w 76 °n 74 °n 72 °n 70 °n 70 °n 72 °n 74 °n 76 °n 50 km 5 8 17 jameson land geographical society ø hold with hope clavering ø kuhn ø store koldewey germania land wollaston forland milne land c4 not resolved pre-devonian rocks fault vertical afta transect afta samples devonian sediments carboniferous–jurassic sediments cretaceous sediments palaeogene basalts cenozoic intrusives cenozoic sediments quaternary sediments ice cretaceous cover o�shore cretaceous cover absent o�shore oceanic crust cretaceous cover prior to c4 exhumation inferred cretaceous cover prior to c4 exhumation fig. 38 outline of the kilometre-thick middle jurassic to upper cretaceous deposits (aalenian–turonian) present prior to exhumation in the mid-cretaceous c4 episode. the map is based on samples of jurassic to cretaceous depositional age that define c4 cooling and older samples that define c4 cooling adjacent to jurassic–cretaceous outliers, e.g. on germania land (appendix 3.13). the cover thus included the post-rift succession deposited after rift-climax at the jurassic–cretaceous transition. due to thermal subsidence, the post-rift succession also covered terrain beyond the rift itself (i.e. ‘steer-head’ geometry). most of this cover was removed prior to palaeogene volcanic extrusions. here, we assume that the extent of the postrift cover reaches c. 50 km away from the data points. we infer that this cover also extended across jameson land where cretaceous sediments are largely absent. c4 cooling could not be resolved in the afta data from jameson land, possibly because of the many later cooling episodes that affected this region. offshore geology according to stoker et al. (2017). http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 67 of 154 www.geusbul let in.org 21°w 21°w vr samples (gc1077) afta samples (gc1077) afta samples vertical transect (gc1016,522) 1 2 3 4 4a 9 10 14 17 16 wollaston forland kuhn ø dombjerg d f pd m f payer land hudson land clavering ø hold with hope episode: palaeotemperature (°c) el ev at io n (k m ) ba se m en t l. c re t. ba se m en t vt16: dombjerg 25°c/km 25°c/km 25°c/km 25°c/km 25°c/km 25°c/km 25°c/km 25°c/km c8 c3 c2c4c5 episode: ba se m en t vt9: kuhn ø c8 c4 c1 episode: ba se m en t pa le oc en e vt17: clavering ø (ne) c8 c2 c1 episode: u pp er c ar bo ni fe ro us vt1: clavering ø (w) c8 c3 c1 episode: pe rm o– tr ia ss ic c8 c4 vt2: kap stosch (hold with hope) episode: ba se m en t pe rm o– tr ia ss ic vt14: giesecke bjerge c8 c6 episode: ba se m en t vt4,4a: hudson land (e) c8 c6 c3 c1 25°c/km 25°c /km episode: u pp er c ar bo ni fe ro us vt3: hudson land (n) c8 c3 c2c6 episode: ba se m en t vt10: payer land c8 c3 c2c9 c4 0 2.0 1.5 1.0 0.5 0.0 20 40 60 80 100 120 140 160 180 palaeotemperature (°c) el ev at io n (k m ) 0 2.0 1.5 1.0 0.5 0.0 20 40 60 80 100 120 140 160 180 palaeotemperature (°c) el ev at io n (k m ) 0 2.0 1.5 1.0 0.5 0.0 20 40 60 80 100 120 140 160 180 palaeotemperature (°c) el ev at io n (k m ) 0 2.0 1.5 1.0 0.5 0.0 20 40 60 80 100 120 140 160 180 palaeotemperature (°c) el ev at io n (k m ) 0 2.0 1.5 1.0 0.5 0.0 20 40 60 80 100 120 140 160 180 palaeotemperature (°c) el ev at io n (k m ) 0 2.0 1.5 1.0 0.5 0.0 20 40 60 80 100 120 140 160 180 palaeotemperature (°c) el ev at io n (k m ) 0 2.0 1.5 1.0 0.5 0.0 20 40 60 80 100 120 140 160 180 palaeotemperature (°c) el ev at io n (k m ) 0 2.0 1.5 1.0 0.5 0.0 20 40 60 80 100 120 140 160 180 palaeotemperature (°c) el ev at io n (k m ) 0 2.0 1.5 1.0 0.5 0.0 20 40 60 80 100 120 140 160 180 25°c /km 30°c /km giesecke bjerge gc1077-70 fig. 39 comparison of palaeotemperature constraints from afta in samples from vertical transects (vts) around clavering ø, plotted against elevation. major offsets are apparent in the magnitude of pre-cenozoic palaeothermal episodes. e.g. between vt10 (in which early jurassic c2 cooling dominates) and vt1 (in which middle triassic c1 cooling dominates). all vertical transects are shown in appendix 2.3. df: dombjerg fault. pdmf: post-devonian main fault. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 68 of 154 www.geusbul let in.org thus, we interpret this episode as representing the time of maximum burial across most of the onshore mesozoic basins between store koldewey and hold with hope and possibly farther south. due to post-rift thermal subsidence, the post-rift succession also covered terrains beyond the rift (i.e. ‘steer-head’ geometry). if we make the conservative assumption that the post-rift cover reached up to 50 km from documented locations, the extent of the former cover would have reached from germania land and southwards to geographical society ø and beyond, and in a wide zone around milne land. it thus seems likely that a thick cover of cretaceous post-rift sediments accumulated above the mesozoic rift basins and their hinterland prior to c4 exhumation. palaeogene basalts are present in the summits across much of the study area (e.g. western hold with hope, kuhn ø and milne land), implying that most of this cretaceous cover had been removed by the time that volcanic activity began. 7.5 fault movements in the clavering ø area palaeotemperature profiles for the clavering ø region in fig. 39 show significant differences between adjacent locations indicating differential exhumation in various events, and major offsets across faults (fig. 16). for example, while results from payer land (vt10) are dominated by the early jurassic c2 episode, those from the nearby clavering ø (vt1) are dominated by the middle triassic c1 episode (appendices 3.8, 3.1). these two locations are on opposite sides of the pdmf, and the results suggest that rocks on the western side (vt10) were more deeply buried in the early jurassic compared to the eastern side (vt1). both locations show earliest cretaceous c4 cooling of broadly similar magnitude, suggesting that the western block was more uplifted during or after the early jurassic episode and that the differential movements were mainly complete prior to the c4 episode. west of the pdmf, the middle triassic c1 episode is recognised in one sample, which records palaeotemperatures in excess of 130°c (gc1077-70). this episode is recognised in many samples east of the fault. this indicates that c1 cooling probably affected areas both west and east of the fault but is only seen in sample gc107770 due to the presence of higher wt% cl apatites, which are more retentive than lower wt% cl grains (wt% cl are provided in supplementary files s1–s3). palaeotemperature profiles for vt9 (kuhn ø) and vt16 (dombjerg) show marked differences for the pre-miocene episodes at these locations (appendices 3.7, 3.10). vt9 cooled below c. 100°c in the middle triassic c1 episode, whereas vt16 cooled below similar temperatures in the early jurassic c2 episode implying that vt9 was colder than vt16 in the early jurassic episode. that relationship persisted to the mid-cretaceous c4 episode when the palaeotemperatures at summit level were about 20°c cooler for vt9 than for vt16. basalts are found in the summits at both locations, so any offsets were levelled out prior to the cenozoic. consequently, the offset between the c2 and c4 palaeotemperatures can be explained by movements of the dombjerg fault during the mid-cretaceous episode, where the dombjerg block underwent more exhumation than the kuhn ø block. at all locations around clavering ø, the summits are at almost identical elevations above present sea level. this implies that any thermal offset between two locations during a palaeothermal episode was levelled out by subsequent differential, vertical movements that ultimately resulted in the uniform present-day temperatures in the summits at similar elevation. offsets of palaeotemperature profiles between two locations can be levelled either by movement of a fault between the two vertical transects (on separate fault blocks; e.g. across the pdmf) or by rotation of the fault block where both transects are located (e.g. vt3 and vt4 on hudson land). basalts occur in the summits of clavering ø, kuhn ø and on dombjerg (vt1, vt17, vt9 and vt16, respectively) at almost the same elevation. this implies that there have been no major relative movements between these locations since the extrusions of the basalts at c. 55 ma. detailed discussion of these differences is beyond the scope of this study, requiring integration with structural information to define different generations of movements. the time-temperature constraints for the different episodes provided in appendix 1.4 allow such investigations by those interested in these aspects of the data. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 69 of 154 www.geusbul let in.org in this chapter, we discuss the hydrocarbon prospectivity of the sedimentary basins in north-east greenland relative to their burial and exhumation history. we investigate when mesozoic sediments may have reached their maximum palaeotemperatures prior to episodes of exhumation and briefly discuss the interplay between igneous bodies, hydrothermal effects, mineralisation and hydrocarbon generation. 8.1 store koldewey afta data from samples of basement and mesozoic sediments in the northernmost part of the study area around store koldewey show that mesozoic sediments reached maximum palaeotemperatures in the mid-cretaceous c4 episode (figs 15, 35, 38; appendix 3.12). for the onshore areas, this implies that hydrocarbon generation from any source rocks would have ceased at the onset of cooling between 95 and 90 ma (c. turonian). successful exploration would therefore depend on longlived seals or remigration. factors such as seal breach during exhumation introduce additional risk. based on the palaeothermal constraints from afta, the basement rocks now exposed on store koldewey were covered by about 2.4 and 2.0 km of sediments prior to mid-cretaceous and late miocene exhumation, respectively (palaeogeothermal gradient 25°c/km, palaeosurface temperature of 20°c). while these conclusions seem to apply across much of the onshore study area (fig. 38), different conditions are to be expected offshore where there was less mid-cretaceous exhumation. source rocks may then have generated hydrocarbons during late cretaceous – cenozoic burial in wide areas across the offshore. exhumation of the margin was also significant during the late eocene and late neogene, but eocene and miocene burial were also important, and the combination of these factors makes it difficult to predict the net effect on thermal history of source rocks some distance from the coast. the interplay between burial and exhumation is likely to exert a critical influence on hydrocarbon prospectivity in offshore basins. correlation between the onshore phases of uplift established in this study and offshore unconformities will be useful in establishing a regional tectonic framework. 8.2 exhumed oil accumulations on traill ø the importance of intrusive activity and exhumation to hydrocarbon exploration in east greenland is illustrated by remnants of exhumed oil accumulations hosted in jurassic sandstones now at elevations of around 1 km on traill ø (price & whitham 1997). the present study provides insight into the magnitude of the sedimentary succession that has been removed on the east greenland margin in the vicinity of these exhumed oil accumulations, together with the timing of the main episodes of uplift and erosion. this provides a framework for understanding the evolution of the margin and associated petroleum systems. in recent years, the implications of these accumulations have been intensely debated (andrews et al. 2020a, b; christiansen et al. 2020). interpretation of afta data from traill ø suggests that prior to the onset of uplift and erosion of the margin in the late eocene, c. 2 km of additional section was present above the present-day summit level of 1 km a.s.l. (section 6.1.1; fig. 26; appendix 3.6). estimates of up to 3 km have been suggested in previous studies (price et al. 1997; price & whitham 1997), but we argue that these studies have underestimated the role of enhanced heat flow associated with the late eocene intrusions in and around eastern traill ø (fig. 32), and their impact on maturity levels, sandstone porosity, permeability and other factors. the generation of hydrocarbons in the vicinity of eastern traill ø from upper jurassic source rocks (price & whitham 1997; price et al. 1997) was most likely enhanced due to the increased heat flow in this region associated with the intrusions in the late eocene (combined with the palaeoburial at that time). the organic content of the reservoir rocks is much lower than would be expected if the petroleum was degraded through biogeochemical processes upon exhumation of the reservoir (j. bojesen-koefoed pers.comm. 2018). the low organic content is thus more likely explained by destruction of previously reservoired hydrocarbons by high temperature, than by leakage and biodegradation. further afield, in areas where heat flow was not enhanced, any source rocks present would have remained at lower maturity levels at the end of the eocene. particularly in offshore regions that were not affected by end-eocene exhumation, such source rocks may then have generated hydrocarbons during later burial in areas of neogene subsidence. therefore, the presence of these exhumed oil accumulations may be seen as a positive feature for exploration in the offshore region, instead of the more negative impact generally attributed to exhumation. 8. hydrocarbon prospectivity http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 70 of 154 www.geusbul let in.org 8.3 maturation studies of the jameson land basin basin modelling and maturation studies of the jameson land basin have used apatite fission-track data to investigate the history of uplift and denudation (christiansen et al. 1992; mathiesen et al. 2000). christiansen et al. (1992) based their exhumation estimate on a range of observations including thermal maturity parameters, apatite fission-track parameters, porosity and seismic velocities; in particular they reported vr values up to 0.6% for organic-rich permian and jurassic formations at outcrop. they found that between 1.5 and 3 km of overburden had been removed across the basin, and that the lost cover likely consisted of more than 1 km of cretaceous sediments as well as 1–2 km of palaeogene basalts (primarily across the southern part of the basin). mathiesen et al. (2000) assumed that generation of hydrocarbons was controlled by the burial resulting from deposition of the thick volcanic pile during the early paleocene and prior to denudation beginning in the late paleocene, with a total of 2–3 km of section removed since that time. in contrast to these earlier studies, our results suggest that early eocene c5 thermal effects in and around the jameson land are dominated not by depth of burial but by hydrothermal and contact effects of igneous intrusive bodies. we find that the regional palaeothermal maximum was reached later, in the end-eocene c6 episode (note, however, that the effects of the mid-cretaceous c4 episode could not be resolved in the afta data from the jameson land basin). because of the localised high c5 palaeotemperatures, severe local variation in maturity levels within potential jurassic source rocks successions is likely to exist in parts of the basin. source rocks will have been heated locally above the gas-generation window. in adjacent areas, similar units may remain immature, while conditions within the main oil-generation window will undoubtedly have been reached in some areas. but amounts of hydrocarbons generated in this episode are likely to be small, given the local extent of early eocene palaeothermal effects. eocene burial following the early eocene c5 episode and prior to maximum cenozoic burial at the regional palaeothermal peak of the c6 episode, will have caused further regional-scale maturation of any source rocks that remained immature after the early eocene. ultimately, the potential yield as well as the potential for migration of hydrocarbons through the succession will, however, be much reduced as a result of early eocene effects (green et al. 2017a). green & japsen (2018) defined three palaeothermal episodes from afta data in two upper jurassic core samples from the 231 m deep blokelv-1 borehole, jameson land (location in fig. 26; east–west profile in fig. 40), combined with vr data and regional afta data. they interpreted localised early eocene palaeotemperatures (corresponding to episode c5) as representing heating related to intrusive activity, and end-eocene c6 and late miocene c8 palaeotemperatures as representing deeper burial followed by successive episodes of exhumation. the end-eocene palaeotemperatures defined by afta require that the marine, upper jurassic succession in the borehole was buried below a 2750 m thick cover of upper jurassic – eocene rocks prior to the onset of end-eocene exhumation (assuming a palaeogeothermal gradient of 30°c/km and 20°c at the surface). as these marine sediments are now near outcrop at c. 200 m above sea level, they have been uplifted by at least 3 km since maximum burial during post-rift thermal subsidence (c. 3.5 km assuming a palaeogeothermal gradient of 25°c/km; fig. 26). the estimated magnitude of rock uplift of about 3 km for sediments penetrated by the blokelv drillcore (green & japsen 2018) is consistent with the results of 3 milne land land surface prior to onset of end-eocene denudation relative to present-day sea level removed section present-day land area w e jameson land liverpool land 25 km blokelv borehole 2 el ev at io n (k m a .s .l. ) 1 0 fig. 8-1 fig. 40 present-day elevation profile from milne land to liverpool land. the elevation of the late eocene land surface relative to present sea level is indicated. end-eocene projected surface elevation (pse) calculated from afta data for jameson land and liverpool land (including results for the blokelv borehole; assuming a palaeogeothermal gradient of 30°c/km) and estimates of rock uplift for milne land (bonow et al. 2014). the late eocene land surface was at low elevation after eocene subsidence, so the elevation of this surface relative to the present sea level is a measure of the amount of rock uplift since then (fig. 26). location of the blokelv borehole and of the profile in fig. 26. modified after green & japsen (2018). http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 71 of 154 www.geusbul let in.org larsen et al. (1989) and bonow et al. (2014). larsen et al. (1989) investigated zeolite isograds in the vesicles of basalts in the region around scoresby sund and argued that the absence of the shallow and less altered zeolite zones could be explained by the removal of these zones by erosion. for milne land they concluded that a succession of about 900 m had been removed above the basalt flows that cover the summits there (assuming a palaeogeothermal gradient of 40°c/km at the time of zeolite formation). bonow et al. (2014) argued that the surface of the volcanic successions was near sea level during the formation of the zeolites because this happened shortly after the eruption of the volcanic units at the paleocene–eocene transition at a time of regional subsidence (brooks 2011). consequently, the basalts, now at c. 1800 m a.s.l., were uplifted 2700 m (1800 + 900 m) since the formation of the zeolites in the eocene. assuming a lower thermal gradient at the time of the zeolite formation would result in a higher estimate of removed rocks. figure 40 shows a present-day elevation-profile across milne land, jameson land and liverpool land and the elevation of the end-eocene land surface relative to present sea level (the end-eocene pse). the late eocene land surface was at low elevation after eocene subsidence, so the elevation of this surface relative to the present sea level is a measure of the amount of rock uplift since then. the influence of younger events in northern jameson land was also recognised in previous studies. temperatures of 115 to 150°c were quoted by christiansen et al. (1990) for formation of various mineral phases in upper permian carbonates in wegener halvø, based on isotope data and fluid inclusion temperatures. christiansen et al. (1992) noted that apatite fission-track ages were very young (25–15 ma) in this region (fig. 12). these observations suggest a clear link between these hydrothermal effects and the early miocene c7 palaeotemperatures revealed by the afta data (fig. 14). this highlights the potential for further studies of the relationship between igneous bodies, hydrothermal effects, mineralisation and hydrocarbon generation in this basin. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 72 of 154 www.geusbul let in.org in this chapter, we compare the regional events of uplift and erosion in greenland and scandinavia. previous studies combining stratigraphic landscape analysis and thermal-history reconstruction in greenland and southern scandinavia have identified a series of post-caledonian regional cooling episodes, mainly representing exhumation (japsen et al. 2006, 2014, 2016, 2018). here we compare the timing of key cooling episodes and identify periods of synchroneity and lack of synchroneity in the onset of cooling across the region (fig. 41). we consider likely mechanisms driving the phases of exhumation and suggest different mechanisms for different tectonic regimes. we discuss the phases of exhumation prior to break-up of pangaea, then the subsequent phases that occurred before the opening of the northeast atlantic. finally, we discuss exhumation in the region after the opening of the north-east atlantic. we also address the development of the modern-day topography and landscapes of greenland and scandinavia. 9.1 phases of exhumation prior to break-up of pangaea in a review of the supercontinent cycle nance et al. (2014) argued that episodic supercontinent assembly and breakup had a profound influence on the course of the earth's geological, climatological, and biological evolution. worsley et al. (1984) argued that supercontinents would become epeirogenically uplifted as heat accumulated beneath them. models of 3d spherical mantle convection have shown that supercontinents shield the underlying mantle from subduction of old oceanic plates, resulting in heat accumulation that produces a 9. episodes of uplift and erosion in greenland and scandinavia u ps lp s lp s u ps formation of peneplain l. neogene (7–2) l. miocene (11–10) 30 0 20 0 10 0 65 60 40 20 0 q central west greenland fig. 9-1 pl mio oli eo pal ku kl j tr pe ca eocene/oligocene (36–30) l. jurassic (160–150) break-up (reburial) (reburial) break-up l. triassic (230–220) e. pliocene (c. 5) l. miocene (c. 10) 30 0 20 0 10 0 65 60 40 20 0 q north-east greenland pl mio oli eo pal ku kl j tr pe ca end-eocene (37–35) e. jurassic (c. 180) m. triassic (c. 240) e. cretaceous (145–140) mid-cretaceous (95–90) c1 c2 c6 c8 c9 c3 c0 c4 l. carboniferous (320–310) l. carboniferous (314–307) h v break-up ti m e (m a) e. pliocene (c. 5) 30 0 20 0 10 0 65 60 40 20 0 q southern scandinavia pl mio oli eo pal ku kl j tr pe ca e. miocene (23–21) m. jurassic (171–164) m. triassic (245–243) mid-cretaceous (107–92) uplift event from stratigraphic landscape analysis time interval during which cooling began onset of cooling (ma) from afta(7–2) fig. 41 timing of regional, post-devonian episodes of uplift and erosion estimated from afta data in north-east greenland (episodes c0–c9; table 1), central west greenland (japsen et al. 2006, 2009) and southern scandinavia (japsen et al. 2007, 2016, 2018). time intervals for the formation of cenozoic peneplains (bonow et al. 2006a, b, 2014; bonow & japsen 2021, this volume; japsen et al. 2006, 2018) are shown. the figure illustrates the similarity of the post-breakup development of the passive margins of west and east greenland, where cenozoic uplift began in the late eocene. this contrasts with southern scandinavia, where cenozoic uplift began in the early miocene, c. 10 myr later than in greenland. reburial of the ups in west and east greenland is tentatively shown at 15 ma. c: carboniferous. dev: devonian. e: early or earliest. eo: eocene. hv: hardangervidda. j: jurassic. kl: lower cretaceous. ku: upper cretaceous. l: late. lps: lower planation surface. m: mid. mio: miocene. oli: oligocene. pal: paleocene. pe: permian. pl: pliocene. q: quaternary. tr: triassic. ups: upper planation surface. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 73 of 154 www.geusbul let in.org concentration of plumes beneath the supercontinent (phillips & bunge 2005, 2007). epeirogenic uplift will dominate tectonic activity during the lifespan of a supercontinent as trapped mantle heat accumulates beneath the largely stationary supercontinent, ultimately manifesting as hot-spot activity and fragmentation (nance et al. 2014). in southern norway, sweden and west greenland, late carboniferous, middle triassic and middle jurassic exhumation episodes have been identified with a timing that is similar or very close to those of the episodes identified in this study (fig. 41; japsen et al. 2016, 2018). these exhumation episodes identified across southern scandinavia were suggested to be the result of the three main phases of epeirogenic uplift accompanying fragmentation of pangaea (japsen et al. 2016). the mid-carboniferous (c. 320 ma) amalgamation of gondwanaland and laurussia and formation of the variscan fold belt resulted in accumulation of heat beneath pangaea and thermal uplift (represented by a stratigraphic gap between 320 and 300 ma) that finally led to the first phase of extension and rifting of pangaea at c. 300 ma (veevers 2004, 2013). the extension generated accommodation space for globally synchronous, cratonic sedimentary successions. the supercontinent was inherently unstable, resulting in almost simultaneous assembly and onset of continental break-up (doré et al. 1999). veevers (2004, 2013) presented a detailed account of the merger and breakup of pangaea based on stratigraphy and radiometric dating. in the following sections we compare their account with results from north-east greenland and elsewhere. 9.1.1 late carboniferous episode (c0) the onset of c0 exhumation in north-east greenland (between 320 and 300 ma) correlates with the onset of a similar period of exhumation in southern scandinavia (between 314 and 307 ma; fig. 41). in north-east greenland, exhumation led to the formation of a peneplain overlain by the upper permian conglomerates of the huledal formation (fig. 18; haller 1971; surlyk et al. 1986; surlyk 1990). in north-west europe, there is evidence of an important phase of exhumation in the southern permian basin towards the end of the carboniferous (doornenbal & stevenson 2010). here, the orogenic activity in the variscan fold belt ceased during the latest westphalian (313–307 ma; middle pennsylvanian). but in large parts of the basin, the stratigraphic record after the cessation of the tectonic activity is represented by a major, latest carboniferous hiatus focused on the stephanian (307–299 ma; late pennsylvanian), corresponding to a pronounced base-permian unconformity. based on the overlap between the timing of late carboniferous exhumation in north-east greenland and north-west europe and stratigraphic constraints, we suggest that the onset of this episode was most likely at c. 310 ma across the entire region. late carboniferous cooling has not been identified as yet in west greenland, but this might be because all rocks studied so far cooled below 110°c in the triassic, so any evidence for carboniferous cooling has been overprinted. veevers (2004) noted that the first stage of extension of pangaea (c. 300 ma) was accompanied by intense magmatic activity in, for example, the danish basin and the oslo rift (vejbæk 1989; larsen et al. 2008b). in east greenland, rifting was not accompanied by volcanism (surlyk 1990). afta studies have also reported late carboniferous exhumation across ireland, namibia, northeast brazil and angola (green et al. 2000, 2009; japsen et al. 2012b; green & machado 2015). we thus consider the late carboniferous phase of exhumation in northeast greenland and southern scandinavia to be part of the first phase of break-up of pangaea. 9.1.2 middle triassic episode (c1) the onset of the c1 episode (c. 240 ma) corresponds closely to that of a similar event in southern scandinavia (beginning between 245 and 243 ma; fig. 41), in scotland (245–225 ma, holford et al. 2010) and in eastern canada (244–225 ma; japsen et al. 2017). we suggest that triassic exhumation in all these areas represents manifestations of a common tectonic event. the uplift and erosion at this time led to the deposition of coarse-grained siliciclastic sediments across the region: the pingo dal group on jameson land (fig. 18; clemmensen 1980; oftedal et al. 2005), the middle–upper triassic skagerrak formation (or hegre formation) around southern scandinavia (bertelsen 1980; goldsmith et al. 2003) and the triassic sherwood sandstone across the british isles. in a review of published triassic stratigraphic nomenclatures for north-west europe, goldsmith et al. (2003) noted that the precise chronostratigraphic equivalence of these units is uncertain. nonetheless, they interpreted deposition of the skagerrak formation to have begun in the anisian at 245 ma. we thus regard the episodes of middle triassic uplift and erosion in north-east greenland and north-west europe as expressions of the same geological event that probably began around 245 ma. the recognition of the middle triassic episode across such a wide area of the northern hemisphere provides further evidence in support of the interpretation of japsen et al. (2016): that this episode represents the second phase of break-up of pangaea. there is a mismatch of 15–20 myr between late triassic cooling in west greenland and middle triassic cooling in other areas (fig. 41). it is unclear if this represents a real diachroneity in cooling across the region or some artefact of the data. however, given the ubiquity of http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 74 of 154 www.geusbul let in.org triassic cooling across a vast area it seems reasonable to suggest that triassic cooling in these areas are expressions of the same phenomenon. the time difference may simply indicate that the episode affected east greenland earlier than west greenland, which would also imply that early mesozoic rifting was probably delayed on the west coast relative to the east coast of greenland. however, this cannot be tested as no triassic sediments have been recorded onshore or offshore west greenland. larsen et al. (2009) found highly alkaline, volatile-rich melts formed in small volumes in the deep lithosphere. these were indicative of incipient stretching in west greenland during late triassic to late jurassic (c. 220–150 ma), and thus indicate delayed triassic rifting in west greenland compared to east greenland. 9.1.3 early jurassic episode (c2) the onset of c2 exhumation (c. 180 ma) identified here is slightly earlier than the middle jurassic event identified from afta data in scandinavia (171–164 ma; fig. 41). surlyk & ineson (2003) showed that the main tectonic events and stratigraphic trends in east greenland, denmark and adjacent areas are highly similar. they found that major regional uplift and erosion took place broadly at the early–middle jurassic boundary (c. 175 ma) in both areas, (fig. 18, 21). subsequent subsidence began in the aalenian (c. 172 ma) in the north sea and in the late bajocian (c. 169 ma) in east greenland associated with the onset of rifting, resulting in major regional onlap and transgression. in north-east greenland, middle jurassic transgression occurred across peneplained basement (surlyk 1977, 1978a, 2003). we thus regard the episodes of uplift and erosion in north-east greenland and north-west europe as expressions of the same geological event that probably began around 175 ma at the early to middle jurassic transition. these observations indicate that the middle jurassic, central north sea dome (ziegler 1990; underhill & parkington 1993; nielsen 2003) is also part of a super-regional phenomenon – the third phase of the fragmentation of pangaea (japsen et al. 2016). afta studies also record mid-jurassic exhumation in northern australia, north-west canada, angola, south africa and north-east brazil where upper jurassic sediments onlap basement (duddy et al. 2004; green et al. 2005; magnavita et al. 2005; green & machado 2015; green et al. 2017b). the mid-jurassic is also a time of copious igneous activity in southern africa (svensen et al. 2012). the timing of the late jurassic cooling episode defined from afta in west greenland (onset between 160 and 150 ma) occurs about 25 myr later than the afta-derived c2 episode in east greenland (fig. 41). no jurassic sediments have been encountered onshore or offshore west greenland, but larsen et al. (2009) found that melts intruded in a 60 km long swarm of scattered alkaline dykes were indicative of increased extension in west greenland around 150 ma (tithonian). the first evidence of rifting west of greenland is dated at 135 ma, much later than in east greenland (chalmers & pulvertaft 2001). in east as well as west greenland, jurassic exhumation led to stripping of the basement over large areas (surlyk 2003; bonow 2005; japsen et al. 2006, 2009). the subsequent transgression began much later in the west where albian sediments rest on basement in the nuussuaq basin, compared to the east, where bajocian sediments rest on basement, for example on wollaston forland (surlyk 2003; dam et al. 2009). we thus conclude that the late jurassic episode of exhumation in west greenland reflects the same geological process as in north-east greenland, but with a considerable delay. in summary, we find it unlikely that exhumation in these disparate locations at three broadly similar times should be a coincidence, and therefore we conclude that exhumation during episodes c0, c1 and c2 recorded in north-east greenland are manifestations of global processes related to the break-up of pangaea. 9.2 phases of exhumation after break-up of pangaea and before the opening of the north-east atlantic 9.2.1 earliest cretaceous episode (c3) the onset of c3 exhumation (145–140 ma) occurs about 20 myr later than the late jurassic event in west greenland (160–150 ma; fig. 41). we consider it likely that these events reflect different geological processes, in particular because the late jurassic episode in west greenland shares characteristics with the early jurassic c2 episode in north-east greenland (section 9.1.3). we interpret the c3 episode as the result of regional uplift and erosion that accompanied the mesozoic rift climax in north-east greenland, not only along the margin of the rift but also of hinterland areas at considerable distance from the rift basins. uplift and erosion associated with the rifting contributed to the coarse-grained basin infill, for example, the wollaston forland group (fig. 18). the c3 episode has no known counterpart based on afta data from scandinavia. however, there are reports of erosion of fault crests and of turbiditic sands at the jurassic–cretaceous transition in the north sea (mcleod et al. 2002; andsbjerg & dybkjær 2003), which may represent a contemporaneous tectonic event. 9.2.2 mid-cretaceous episode (c4) the c4 episode (beginning between 95 and 90 ma; fig. 41) most likely corresponds to the mid-turonian, erosional base of the jackson ø group that records a phase of basin reorganization (fig. 18; bjerager et al. 2020). we suggest that this episode represents regional http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 75 of 154 www.geusbul let in.org exhumation that led to removal of much of post-rift succession that had accumulated above the rift and its margins in the interior of north-east greenland after the rift climax at the jurassic–cretaceous transition (fig. 38; sections 5.5 and 7.4.4). the lofoten margin of northern norway forms the conjugate margin to north-east greenland. here, an episode of early late cretaceous uplift and erosion (beginning in the middle cenomanian; c. 95 ma) is represented by a regional unconformity (henstra et al. 2016b). further south on the møre-trøndelag margin of western norway, a significant intra-turonian unconformity most likely formed in response to a regional, extra-basinal process (sømme et al. 2013). these observations match the c4 episode in north-east greenland, suggesting a common tectonic phase. the c4 episode is coeval with the onset of inversion of the sorgenfrei–tornquist zone along the southern margin of the baltic shield and of the sole pit axis in the southern north sea, and a regional turonian unconformity in the faroe–shetland basin (mogensen & korstgård 2003; japsen et al. 2007, 2016; stoker & ziska 2011; green et al. 2017c). kley & voigt (2008) argued that a late cretaceous event of intraplate basin inversion and basement thrusting in central europe occurred after a change in relative motion between the european and african plates at c. 90 ma. the pinching of west-central europe’s thin lithosphere between baltica and africa was attributed to this contraction. these events may be related to a global-scale, mid-cretaceous phase of plate reorganisation (105–100 ma; matthews et al. 2012). there is no known counterpart for the c4 episode in west greenland. 9.3 phases of exhumation after opening of the north-east atlantic 9.3.1 end-eocene episode (c6) the onset of c6 exhumation in east greenland is contemporaneous with exhumation in west greenland that began at the eocene–oligocene transition and thus affected both margins at about 35 ma (figs 23, 41). these events led to the formation of regional peneplains on both sides of greenland (both labelled ups; bonow et al. 2006a, b, 2014; bonow & japsen 2021, this volume), and we conclude that they reflect the same geological phenomenon. this episode also correlates with renewed magmatic activity in east greenland, together with block rotation and faulting (figs 28, 29). offshore north-east greenland, a significant phase of progradation occurred above a horizon that defines an erosional incision, dated broadly from late eocene to mid-miocene (petersen 2019; fig. 6e). the clinoforms are caused by extensive mass wasting into the outer part of the shelf due to tectonic uplift of the inner part of the shelf and onshore areas, south of store koldewey. we suggest that the mass wasting is due to the uplift of the margin of north-east greenland during the c6 episode. uplift of the inner margin of south-east greenland resulted in strong influx on the shelf of coarse, siliciclastic turbidites above a middle eocene to upper oligocene unconformity (larsen et al. 1994a; japsen et al. 2014). oligocene sediments are absent on the shelf off central west greenland, and the hiatus is represented by a low-angle unconformity on seismic sections (sørensen 2006). green et al. (2013) correlated the oligocene un con formity offshore with the ups onshore. late eocene exhumation also affected the barents sea, svalbard and the north slope of alaska according to afta data (green & duddy 2010). late eocene exhumation coincided with the cessation of sea-floor spreading west of greenland and with cessation of spreading on aegir ridge, and the establishment of the kolbeinsey ridge east of greenland (gaina et al. 2009). plate reorganisation was preceded by a sharp change in the direction of spreading and a drop in spreading rates, such that the eurasian and greenland plates were almost stationary relative to one another (gaina et al. 2009, 2017). whatever the causes of these phenomena, they indicate substantial changes in the stress regime in and around the north-east atlantic at the eocene–oligocene transition (japsen et al. 2014). interestingly, afta data from southern scandinavia do not reveal a counterpart to episode c6 (fig. 41). however, a regional, intra-upper eocene unconformity extend along the north-west european margin, where a dramatic change in basin configuration in the region took place during late eocene – early oligocene (stoker et al. 2005). this shows that major tectonic events at this time affected wide areas in the north-east atlantic domain and beyond. 9.3.2 early miocene uplift and erosion in southern scandinavia neogene uplift and erosion in southern scandinavia began in the early miocene (between 23 and 21 ma; fig. 41) and resulted in the formation of a regional peneplain, which today is represented by the hardangervidda plateau in southern norway and the south småland peneplain in southern sweden (japsen et al. 2016, 2018). the erosion in the scandinavian hinterland resulted in the formation of braided fluvial systems, progradation of lower miocene deltas and deposition of thick, coarse-grained successions across much of the present-day danish onshore area (rasmussen et al. 2010; rasmussen 2014). a major, base-neogene unconformity developed in the sedimentary basins along the margin of north-west europe (stoker et al. 2005). the early miocene uplift event has no known counterpart in greenland. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 76 of 154 www.geusbul let in.org 9.3.3 late miocene episode (c8) exhumation in the c8 episode in east greenland began at c. 10 ma and also affected the west greenland margin (figs 23, 41). on both margins, this event initiated the formation of the present relief with incision below the uplifted ups leading to the development of the lps. it seems likely that a common process was responsible on both margins. døssing et al. (2016) demonstrated that the onset of mid-late miocene uplift of the north-east greenland led to formation of a regional imu and to massive progradation on the shelf (fig. 6). correlation between margin uplift and plate motion changes indicated that the uplift was triggered by plate tectonic forces. we suggest that the uplift onshore and the subsequent progradation offshore began during the c8 episode. late neogene progradation along the east greenland margin has been documented previously (clausen 1998; larsen et al. 1994a; hamann et al. 2005; berger & jokat 2008, 2009; petersen 2019). the c8 episode has no known counterpart in southern scandinavia, but green & duddy (2010) defined a regional phase of exhumation of svalbard and the barents sea beginning between 10 and 0 ma from afta data, in agreement with results from svalbard (dörr et al. 2012, 2019). 9.3.4 early pliocene episode (c9) the timing of uplift and incision in north-east greenland in the c9 episode (c. 5 ma) is within the limits for onset of latest miocene-pliocene uplift defined from afta data in west greenland (7–2 ma; fig. 41). on both margins of greenland, the ups and lps were uplifted and the modern relief was formed by incision of valleys and fjords below the lps. we thus consider it likely that the same geological process was responsible for these events on both margins. in scandinavia, the hardangervidda plateau and the south småland peneplain reached their present elevations of 1200 and 150 m a.s.l., respectively, due to early pliocene uplift (japsen et al. 2016, 2018). neogene exhumation of the danish basin began at c. 5 ma as revealed by afta data, and a regional, intra-neogene (c. 4 ma) unconformity across north-west europe developed in response to pre-glacial uplift and tilting of the continental margin (stoker et al. 2005; japsen et al. 2007). it is therefore possible that early pliocene uplift affected west and east greenland and southern scandinavia within a very narrow time interval, and that the uplifts may reflect the same geological process. the much higher elevation of east greenland compared to west greenland may indicate dynamic support in the east from the iceland plume, particularly during the c9 episode that lifted the lps to its maximum elevation of 2 km a.s.l. in the blosseville kyst region – host to the highest peaks in greenland (japsen et al. 2014). this assertion matches the results of steinberger et al. (2015) who used a mantle-flow model to show that dynamic topography reaches a maximum of 1.2 km for greenland along blosseville kyst. 9.4 development of the modern topography in greenland and scandinavia in chapter 6, we presented a model for the development of the modern topography in north-east greenland that described three episodes of uplift and erosion (fig. 34). in central west greenland and in southern east greenland, the mountains were formed by a similar sequence of events, and the correlation of these episodes in fig. 41 highlights that these uplifts are simultaneous in west and east greenland within the resolution of the methods applied (japsen et al. 2006, 2014). we conclude that the present topography of west and east greenland were formed almost simultaneously in the late neogene and by the same processes. in contrast, the mountains of southernmost norway formed during two steps (japsen et al. 2018). first, early miocene uplift and erosion to the base level defined by the adjacent ocean led to formation of a peneplain that included the present-day hardangervidda plateau and the south småland peneplain (lidmar-bergström et al. 2000, 2013). second, early pliocene uplift raised these peneplains to their present elevations and led to re-exposure of a tilted, mesozoic surface at lower elevations. it is thus possible that the final phase of uplift, in the early pliocene, affected both greenland and scandinavia at about the same time, and thus that this phase was caused by the same processes, of which dynamic support from the iceland plume may be one component. otherwise, the previous cenozoic episodes of uplift of the landscape are out of phase across the atlantic ocean (fig. 41). in greenland, the formation of the dominant peneplain began after uplift in the late eocene, whereas the formation of the dominant peneplain in southern norway began in the early miocene. uplift of these plains, which led to the formation of the present-day relief, began in the late miocene in greenland and in the early pliocene in scandinavia. as noted by japsen et al. (2014), synchroneity be tween changes in plate motion and uplift events along conjugate margins is suggestive of rapid, far-field transmission of momentum through the upper mantle by viscous stresses (colli et al. 2018). conversely, uplift events that have no counterpart on the conjugate margin may be related to transmission of stress within the crust (cloetingh & burov 2010). the correlation – and lack of correlation – between uplift events across the north-east atlantic (fig. 41), thus invites speculation about the underlying processes: http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 77 of 154 www.geusbul let in.org 1. end-eocene exhumation affected arctic regions. this episode coincides with the cessation of sea-floor spreading west of greenland and with a major plate reorganisation in the north-east atlantic, preceded by a significant reduction in spreading rates (gaina et al. 2009, 2017; green & duddy 2010). considering the location of the region affected, it may not be surprising that the effects in southern scandinavia were minor (japsen et al. 2018). the forces driving the uplift are therefore most likely linked to the stress that prevented further movements of the greenland plate relative to north america. 2. regional early miocene uplift only affected the european side of the north-east atlantic. this basic observation led japsen et al. (2016) to suggest that the forces driving uplift most likely originated from stress transmitted from orogenic belts elsewhere on the eurasian plate (cloetingh & burov 2010). a possible source for such stress could be the early miocene–present neohimalayan orogenic phase in himalaya and southern tibet (hodges 2000; jiang & li 2014). the ‘hard’ india– asia collision around 25–20 ma initiated this phase and coincides with deformation in central asia (van hinsbergen et al. 2012). an alternative source could be the collision between africa and eurasia. demets et al. (2015) estimated that nubia-eurasia plate motion decreased by c. 50% between 20–18 ma and 13 ma. 3. major late miocene uplift affected west and east greenland as well as svalbard and the barents sea but not southern scandinavia. the uplift was triggered by plate tectonic forces (døssing et al. 2016), and we speculate that late neogene changes in the absolute motion of the north american plate may explain this asymmetric behaviour across the north-east atlantic. sea-floor spreading rates between eurasia – north america and nubia – north america varied modestly from 20 to c. 8 ma but slowed down by c. 20% between 8 and 5 ma. rates have remained remarkably steady since then (merkouriev & demets 2008; demets et al. 2015; iaffaldano & demets 2016). these changes in relative motions were related to a change in the absolute motion of north america and antarctica, and they might be linked to the late neogene dynamics of the pacific plate (iaffaldano & demets 2016). 4. early pliocene uplift affected west and east greenland as well as southern scandinavia. in scandinavia, the magnitude of rock uplift is about 1 km in southern norway but only 150 m in southern sweden (lidmar-bergström et al. 2013; japsen et al. 2016, 2018). in west and east greenland, it is typically 1 km and in central greenland it is likely to be negligible (solgaard et al. 2013). however, along blosseville kyst, where greenland is closest to iceland, it is about 2 km, possibly indicative of support from the iceland plume (rickers et al. 2013; bonow et al. 2014; japsen et al. 2014; steinberger et al. 2015). the fact that margins on both sides of the north-east atlantic were affected by early pliocene uplift, suggests that this episode may have been driven by changes in upper-mantle flow, related to the iceland plume (colli et al. 2018). http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 78 of 154 www.geusbul let in.org in this chapter, we illustrate the importance of imposing geological constraints on the interpretation of thermal data and discuss the role of the neogene uplift for the development of the greenland ice sheet. finally, we place the development of the east greenland margin in the context of epcms around the world. 10.1 imposing geological constraints on thermal-history reconstructions the development of the continental margin of east greenland from 68°n to 78°n described here and by japsen et al. (2014) is very different to previous ideas for the development of epcms in general (see green et al. 2013) and east greenland in particular (chapters 3 and 4). one key reason for this difference is that whereas our studies have integrated thermal histories from afta with geological and landscape evidence, other studies largely ignore the abundant geological constraints exposed in north-east greenland. for example, johnson & gallagher (2000) reported thermal histories derived from apatite fission-track data in samples of carboniferous sandstones from clavering ø. however, they did not consider that palaeogene basalts cap the sampled profile, and thus that the uppermost sample was at near-surface temperature in the palaeogene. this led them to underestimate the final degree of mesozoic cooling and palaeogene re-burial prior to the onset of cenozoic cooling (japsen et al. 2010; green et al. 2013). we present a thermal-history diagram based on our interpretation of afta data from clavering ø (vt1) in appendix 3.1. some authors have argued that the mountains in north-east greenland and in scandinavia are remnants of the caledonian mountains (nielsen et al. 2009; pedersen et al. 2012; schiffer et al. 2016). however, the highest mountain in greenland (gunnbjørn fjeld) is comprised of paleocene basalt, and all of the high terrain in east greenland south of 70°n, as well as all of west greenland, is outside the caledonian orogen. in addition, sediments of different age rest on the basement within the former orogen where mid-permian, middle jurassic and younger peneplains also occur (haller 1971; larsen 1988; surlyk 2003; bonow & japsen 2021, this volume). these observations testify to a history of episodic exhumation and burial rather than to a history of continuous exhumation since the caledonian orogeny. pedersen et al. (2012) suggested that the present-day mountains in north-east greenland represent remnants of the original caledonian topography modified during ?devonian – early carboniferous rifting, and that the area has undergone slow, continuous exhumation since c. 250 ma. but they did not take into account simple, yet robust geological constraints on the post-caledonian development of the region (fig. 42). pedersen et al. (2012) based their hypothesis on inverse modelling calibrated by new apatite fission-track data from samples of exposed caledonian basement. however, they did not consider that apatite fission-track data only register episodes of cooling from higher temperatures. further, they did not consider that the presence of carboniferous and jurassic sedimentary remnants across their study area define times when the underlying basement was at the surface prior to a phase of burial and heating. thus, integrating the apatite fission-track data with geological evidence defines multiple episodes of exhumation (cooling) and re-burial (heating). the well-documented geological record of north-east greenland shows that the caledonian mountains were obliterated as topographic features during the late palaeozoic, and the post-caledonian record provides clear evidence of episodic, exhumation and re-burial (haller 1971; surlyk 1990; japsen et al. 2013). the episodic nature of the development of germania land after the caledonian collapse (points 1 to 7 in fig. 42), is as follows: 1. caledonian basement with eclogitic inclusions reflects high-pressure metamorphism of devonian age (700– 800°c, 410–390 ma; gilotti & mcclelland 2008). 2. upper carboniferous sediments resting on the caledonian basement shows that the basement was exhumed to the surface by the late carboniferous (piasecki et al. 1994). 3. middle triassic c1 palaeotemperatures of 100°c from afta indicate that these carboniferous units were buried below a 3.2 km thick cover of upper carboniferous to middle triassic sediments in the middle triassic (samples gc1077-9, 10, 11; map c1 in fig. 14; appendix 3.13; palaeogeothermal gradient 25°c/km, 20°c surface temperature). 4. the presence of middle jurassic sediments overlying the basement shows that the upper carboniferous – middle triassic cover was removed by the middle jurassic (bojesen-koefoed et al. 2012). 5. mid-cretaceous c4 palaeotemperatures of c. 75°c from afta indicate that the jurassic units were buried below a 2.2 km thick cover of middle jurassic to turonian sediments in the mid-cretaceous (samples gc1077-9, 11; thermal parameters as above). 6. late miocene c8 palaeotemperatures of c. 65°c from afta indicate that these deposits were buried below a 10. discussion http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 79 of 154 www.geusbul let in.org representative sample and sample number aft sample, pedersen et al. (2012) 20 20 30 28 vr c. 0.5% vr c. 0.5% vr c. 0.5% palaeogene sediments palaeogene basalts middle jurassic – cretaceous sediments upper carboniferous – permian sediments caledonian basements ice pre-caledonian rocks sedimentary outlier 100 km store koldewey kulhøj depotnæsset germania land kuhn ø clavering ø hochstetter forland holm land 12°w 12°w 20°w 20°w 74°n 76°n 78°n 80°n 0 ord si dev carb pe tr jur cret cen 1 2 3 4 5 0100200300 time (ma) 400500 group 1 sog20 exhumation and cooling histories from pedersen et al. (2012) >3 k m d is cr ep an cy be tw ee n m od el a nd d at a d ep th (k m ) a 2 4 7 b c d 2 4 1 3 4 5 6 2 1: caledonian metamorphism; 700–800ºc 2: exhumed to surface by late carboniferous 3: buried by 3.5 km of u. carboniferous to m. triassic sediments 4: exhumed to surface by middle jurassic 5: buried by 2.5 km of middle jurassic to turonian sediments 6: buried by 2 km of middle jurassic to miocene sediments rocks exhumed to the surface 7: exhumed to surface at present day fig. 42 conflicting interpretations of the post-caledonian development in north-east greenland. a: geology of north-east greenland and location of apatite fission-track samples analysed by pedersen et al. (2012). sedimentary outliers occur between 75 and figure 42 continued on next page http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 80 of 154 www.geusbul let in.org 1.8 km thick cover of middle jurassic to miocene sediments in the late miocene (several samples; thermal parameters in point 3). 7. today the caledonian basement is exposed at the surface over most of the region. the interpretation of apatite fission-track data in terms of monotonic cooling as presented by pedersen et al. (2012) is thus inconsistent with the presence of the sedimentary outliers and fails to capture the episodic nature of the post-caledonian history. 10.2 tectonic uplift vs erosional unloading geological and topographical observations from the scoresby sund area showed that erosional unloading due to incision below the highest peaks in the area can account for 1.1 km of uplift, comparable with the elevation of mesozoic marine sediments now at 1.2 km a.s.l. (medvedev et al. 2008). later modelling results of all greenland found that uplift due to erosional unloading could only account for 300–600 m of present elevation of marine sediments (medvedev et al. 2013). medvedev et al. (2013) quantified the influence of the load of massive ice sheets on the greenland topography and the significant erosion caused mainly by glacial carving. they concluded: (1) that the ice-sheet load causes up to 850 m subsidence of the bedrock topography in central greenland, whereas the peripheral bulging caused by this ice loading has a negligible effect on uplift of the greenland margins. (2) that glacial carving can explain up to 1.2 km uplift along the central parts of the margins of east and west greenland. (3) that ice-related processes cannot explain much of greenland’s topography, and that the origin of the pre-glacial mountains was enigmatic. medvedev et al. (2008, 2013) did not consider the possibility that a sedimentary pile of unknown thickness may have accumulated above mesozoic sediments in the scoresby sund area. if it had, the total amount of rock uplift of the exposed sediments since their maximum burial may have been much larger than 1.2 km. for example, afta data from samples of jurassic sediments within the study area of medvedev et al. (2008), show that these samples, which are now at an elevation of 1200 m a.s.l., were heated to between 65 and 85°c prior to end-eocene c6 cooling (vt13, northern jameson land; appendix 2.3). such palaeotemperatures correspond to km-scale burial of the jurassic rock samples for any reasonable palaeogeothermal gradient. thus, uplift of jurassic sediments since end-eocene maximum burial is likely to be several times larger than the maximum amount allowed by the calculations of medvedev et al. (2008). the ups forms the high ground west and south of scoresby sund. as an oligocene–miocene peneplain formed by erosion to sea level, the fully developed peneplain provided no relief to be eroded at the onset of uplift in the late miocene. thus, even if the erosion rate increased due to climate deterioration, no valley incision could have occurred to initiate the uplift. as there was no significant change in eustatic sea level at this time (miller et al. 2005), a tectonic trigger was required (japsen et al. 2014). 10.3 development of the greenland ice sheet 10.3.1 late neogene onset of the formation the greenland ice sheet using an ice-sheet model coupled with a climate model, solgaard et al. (2013) studied the build-up of ice caps before and after late miocene c8 and early pliocene c9 uplifts (fig. 34). they applied forcing conditions for the late cenozoic climate (i.e. long-term cooling superimposed by cold and warm climate excursions) and found that no ice covers were initiated for the low-lying and almost flat topography prior to late miocene c8 uplift (i.e. the ups at sea level). they also demonstrated a significant ice-sheet growth in response to the orographically induced increase in precipitation and the cooling of surface temperatures accompanying uplift. large amounts of ice formed after the first (c8) uplift of the ups, but the modelling showed that the ice sheet was sensitive to changes in climate. the second (c9) uplift, however, facilitated ice-sheet build-up and increased the stability of the ice sheet by providing anchoring points that were not available to the same extent in the previous scenarios. however, a föhn effect was also observed that inhibited ice-sheet expansion into greenland’s interior and figure 42 continued 80°n and therefore the underlying basement was at the surface prior to their deposition at these locations. corresponding sedimentary units are more extensive farther south and north. grey circles: samples sog20, 28, 30 from pedersen et al. (2012), chosen as representative of their three groups of samples. b: outcrop of middle jurassic coal (triangle 4 in a). c: upper carboniferous deposit (triangle 2 in a). fossil and reconstruction of lepidodendron (piasecki et al. 1994). d: exhumation path (red) based on modelling by pedersen et al. (2012) of apatite fission-track data in a representative sample chosen by these authors. yellow curve: constraints from the geological record and afta data on the burial and exhumation history of basement rocks on germania land (see appendix 3.13). ord: ordovician. si: silurian. dev: devonian. carb: carboniferous. pe: permian. tr: triassic. jur: jurassic. cret: cretaceous. cen: cenozoic. modified after japsen et al. (2013) and green et al. (2013). http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 81 of 154 www.geusbul let in.org thus shifted the threshold of formation of the ice sheet towards colder temperatures. under conditions that are colder than the present, the ice can overcome the föhn effect, flow into the interior and form a coherent ice sheet. the results thus indicated that the greenland ice sheet of today is a relict formed under colder conditions. these modelling results are consistent with the observed climatic history in greenland. late miocene onset of uplift of the east greenland margin correlate with the onset of neogene progradation, north atlantic bottom water and siliciclastic neogene sedimentation at c. 10 ma and the first occurrence of glacial sediments in drill cores off south-east greenland at c. 7 ma (fig. 6; fredskild et al. 1989; larsen et al. 1994b; berger & jokat 2008, 2009; døssing et al. 2016). the onset of miocene progradation off north-east greenland therefore reflects the uplift of the margin and the subsequent incision below the ups. the supply of glacial deposits to the shelf reflects the formation of glaciers on the newly-formed mountains onshore. modelling of the greenland ice sheet shows that glaciers could not have formed in the miocene prior to the uplift in episode c8 because of the low absolute relief at that time (solgaard et al. 2013). thus, the late cenozoic mountain building in greenland augmented the effects of the climatic deterioration leading to the northern hemisphere glaciations. without the pliocene phase of uplift the greenland ice sheet would have been more sensitive to the changes in climate over the past millions of years (solgaard et al. 2013). 10.3.2 continental ice in east greenland at the eocene–oligocene transition several studies have suggested that continental ice may have developed in east greenland at the eocene–oligocene transition. stratigraphically extensive, ice-rafted debris, including macroscopic dropstones, occur in late eocene to early oligocene sediments from the norwegian–greenland sea, indicating sediment rafting by continental ice, not sea ice, and east greenland as the likely source (eldrett et al. 2007). eldrett et al. (2009) presented climate estimates for the eocene to oligocene interval, based on spore and pollen assemblages in marine sediments from the norwegian–greenland sea. the climate estimates indicated cooling across the eocene–oligocene transition, but also provided evidence for relatively warm summer temperatures at that time, and thus that continental ice on east greenland was probably restricted to alpine outlet glaciers. the evidence presented here suggests that the appearance of continental ice in the late eocene may correspond to the onset of the regional, end-eocene c6 phase of uplift of the east greenland margin. the late eocene land surface in east greenland was likely to be of low relative relief and near the level of the adjacent sea after eocene post-rift subsidence and burial. the onset of exhumation by the end of the eocene is thus likely to imply that the surface was uplifted and that a terrain of some elevation must have formed. even if the altitude of the terrain at the eocene–oligocene transition in southern east greenland is unknown, it is likely to have been significant because a rock column up to 2 km was removed between the late eocene and the late miocene in the kangerlussuaq area, south of blosseville kyst (japsen et al. 2014). we thus suggest that the high terrain generated by the end-eocene uplift was sufficient to trigger the formation of continental ice in east greenland. we also interpret the lack of indications of continental ice along the east greenland margin between the late oligocene and the late miocene to reflect the absence of high mountains during the bevelling of the landscape and the formation of the ups by erosion to base level prior to uplift in the late miocene c8 episode. a phase of exhumation starting at 30 ± 5 ma has been identified in thermochronology data from north-east greenland (bernard et al. 2016). this is consistent with uplift and erosion in the end-eocene c6 episode defined here. however, the authors explained the exhumation as the result of glacial erosion initiated by a worldwide temperature drop at the eocene–oligocene transition (zachos et al. 2001). this interpretation contradicts several observations. first, continental ice in east greenland at that time was probably restricted to alpine outlet glaciers (eldrett et al. 2009) and thus could not produce the regional, end-eocene exhumation documented in this study (fig. 14). second, end-eocene exhumation involved block rotation and faulting, observations that are not be related to climatic conditions (figs 28, 29). third, the end-eocene phase of exhumation in greenland coincides with magmatic activity in east greenland (fig. 21), cessation of sea-floor spreading west of greenland and a major plate reorganisation in the north-east atlantic (gaina et al. 2009), and therefore must have a tectonic origin. 10.4 episodic tectonic processes along margins and their hinterland north-east greenland is a typical example of an epcm where the topography developed long after break-up (japsen et al. 2012a; green et al. 2013, 2018). the distinct plateaus, the ups and lps, define the large-scale landscape of north-east greenland and are visual evidence for episodic tectonic processes that shaped the margin after break-up at the paleocene–eocene transition. in general, elevated plateaus are evidence of at least two major phases of uplift. an initial phase produces the peneplain as an erosion surface that cuts http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 82 of 154 www.geusbul let in.org unconformably across rocks of different age and resistance, and a second phase of uplift brings the peneplain to its present elevation above sea level. formation of such elevated plateaus bordering mesozoic–palaeogene rift basins as a result of neogene uplift, represents a major challenge to conventional geological thinking, which predicts post-rift thermal subsidence and deposition of a thick post-rift succession overlying both the rift and its margins (mckenzie 1978; white & mckenzie 1988). however, numerous studies have documented significant neogene uplift of many epcms, not only of margins in the north-east atlantic domain but also elsewhere, in different climatic regimes (japsen et al. 2012b; green et al. 2013; richards et al. 2016; rodríguez et al. 2017; da silva et al. 2019). the development of the east greenland margin was also characterised by episodic processes prior to break-up as it is well documented in numerous studies of the post-caledonian geology (surlyk 1978a, 2003; clemmensen 1980; stemmerik 2000; larsen et al. 2008a). in particular, three phases of uplift and denudation that affected north-east greenland and scandinavia in the late carboniferous, middle triassic and early–middle jurassic, appear to be part of widespread epeirogenic uplifts that accompanied fragmentation of pangaea (section 9.1). along epcms elsewhere, episodic burial and exhumation, rather than slow monotonic denudation, affected the sedimentary basins and their hinterland both prior to and after break-up (green et al. 2013, 2018). most thermochronological studies of basement terrains are carried out within a framework of continuous cooling. however, many of the maps in fig. 14, from basement in the west to sedimentary basins in the east, show little variation in palaeotemperatures of individual episodes, suggesting a similar, shared thermal history. this is supported by the close correlation between, for example, the middle triassic uplift and erosion of the basement terrains in the interior of north-east greenland and the deposition of the coarse clastic material of the scoresby land group in the rift basin. thus, episodic cooling and exhumation also affected the hinterland far from the basin and was likely a plate-tectonic-scale event. the presence of palaeozoic, mesozoic and palaeogene rocks on basement in the study area further underlines the episodic nature of the burial and exhumation history. we therefore conclude that episodic heating and cooling is the most appropriate framework for extracting thermal-history solutions from thermochronological data, both in the sedimentary basins and the basement terrains of north-east greenland, as documented in other parts of the world. deciphering past geological events and modelling the physical processes that drive them are fundamental for our understanding of geological processes, such as sedimentary basin formation and the uplift of mountain belts. efforts to date have centred largely on continuous, long-term subsidence or uplift, but here, we document that the geological record is very incomplete, and histories typically involve repeated episodes of upward and downward vertical movement. a detailed understanding of the nature of the processes underlying the development of epcms (as well as other geological features) requires improved geodynamic modelling to develop and test likely mechanisms (green et al. 2018). the observations for north-east greenland presented here can provide constraints against which projections from geodynamic models can be tested. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 83 of 154 www.geusbul let in.org 11.1 palaeothermal episodes our results reveal a long history of episodic burial and exhumation across north-east greenland, where denudation across present-day basement terrains in the interior produced sedimentary input to the palaeozoic– mesozoic rift basins. the earliest evidence for cooling in the afta data defines the late carboniferous c0 episode (that began between 320 and 300 ma) of uplift and erosion that led to formation of the sub-permian peneplain. today, this peneplain is exhumed on jameson land from below upper permian cover rocks, for example the huledal formation. the corresponding unconformity marks the most profound change in tectonic history and depositional environment in the post-caledonian development of east greenland. late carboniferous uplift and erosion also affected north-west europe, and we suggest that the onset of this episode in both regions most likely happened around 310 ma. in many areas, palaeozoic and older rocks were buried below a cover with a thickness of 3 km or more prior to exhumation in the middle triassic c1 episode (c. 240 ma). in the region from germania land to clavering ø, the cover consisted of upper carboniferous – middle triassic rocks. regional uplift and erosion in the middle triassic produced thick, siliciclastic deposits in the rift system, such as the pingo dal group. middle triassic uplift and erosion also affected north-west europe at about the same time, and we suggest that the onset of this episode in both regions probably happened around 245 ma. early jurassic c2 exhumation (c. 180 ma) caused denudation of the basement across much of the study area apart from the jameson land basin. the episode is reflected in the base-middle jurassic unconformity and in an etch (weathering) surface above which middle and upper jurassic sediments accumulated, for example charcot bugt, pelion and payer dal formations. previous studies interpreted the early jurassic uplift and erosion to be caused by a dome centred north of jameson land, but our results show that the exhumation also affected areas far into the hinterland. uplift and erosion also affected north-west europe at about the same time, and we suggest that the onset of this episode in both regions probably began around 175 ma, at the early to middle jurassic transition. earliest cretaceous c3 exhumation (145–140 ma) accompanied the late mesozoic rift climax north of kong oscar fjord, and the subsequent erosion contributed to the coarse-grained sedimentary infill of the rift basins (the wollaston forland group). these basins were subsequently buried during post-rift thermal subsidence. sporadic evidence for episode c3 in the south-west indicates that uplift most likely affected much of the region. uplift and erosion that began in the mid-cretaceous c4 episode (95–90 ma) affected large parts of the region and corresponds to the mid-turonian, erosional base of the jackson ø group. when exhumation began, a cover of cretaceous sediments, up to 2 km thick, including post-rift sediments, had accumulated above the mesozoic rifts and their hinterlands. where the c4 episode represents maximum palaeotemperatures for the mesozoic sediments, it thus also marks the termination of hydrocarbon generation (between store koldewey and hold with hope and possibly further south). the effects of this exhumation episode could not be resolved in the afta data from the jameson land basin, where cretaceous sediments are largely absent, possibly due to the many later cooling events that affected this region. fault movements of up to 1 km affected the clavering ø region in this episode, but the offsets were levelled out and an etch surface developed prior to eruption of the palaeogene basalts during breakup. our results show no evidence for regional exhumation related to the final rifting and breakup at the palaeocene–eocene transition. this agrees with geological evidence that demonstrates continued subsidence and burial during and after break-up. we interpret early eocene c5 cooling (c. 55 ma) focussed around jameson land to reflect local heating associated with widespread eocene intrusive activity. following break-up, a cover of palaeogene basalts and eocene sediments accumulated on the rifted margin and began to be exhumed during end-eocene c6 uplift (37–35 ma). when uplift began, all rocks that are now exposed at the surface in a zone from wollaston forland to liverpool land, were buried below a considerable rock column, typically a cover with a thickness of 1 km or more. the end-eocene episode affected the entire east greenland margin. this episode coincided with the cessation of sea-floor spreading west of greenland, a major plate reorganisation in the northeast atlantic and renewed intrusion of magmatic bodies along the coast, notably on traill ø where palaeotemperatures were elevated locally by heat flow close to the intrusions. the early miocene cooling c7 episode (20–18 ma) was focussed on northern jameson land. it may have involved local uplift and erosion, possibly related to the oligocene intrusive centre of werner bjerge and 11. conclusions http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 84 of 154 www.geusbul let in.org malmbjerget, and the formation of the high terrain of stauning alper at the western edge of these intrusive centres. uplift in the late miocene c8 episode (c. 10 ma) initiated the development of the modern topography in east greenland. the present relief of fjords and valleys across the entire region were incised after the late miocene. in the interior highlands, the incision took place below the general summit level in these mountains. in the coastal zone, the present-day summits were buried below a significant rock column of primarily miocene sediments prior to late miocene uplift and erosion. the early pliocene c9 episode (c. 5 ma) is primarily defined by afta data in the region around the intrusive centre on traill ø. here, end-eocene palaeotemperatures are high due to heat from the intrusive bodies. early pliocene palaeotemperatures are high due to deep burial. 11.2 development of present-day topography two elevated peneplains extending across the study area provide visible evidence of three cenozoic phases of uplift and erosion defined from afta: (1) uplift during the end-eocene c6 episode led to formation of the ups by erosion to the base level defined by the adjacent atlantic ocean. (2) uplift during the late miocene c8 episode of c. 1 km led to formation of the lps by incision along rivers below the uplifted ups towards the new base level. (3) uplift during the early pliocene c9 episode of c. 1 km lifted the ups and the lps to their present elevations of c. 2 and 1 km a.s.l., respectively, and initiated the formation of the present-day landscape through fluvial and glacial erosion. this implies that the low-lying ups dominated the landscape during most of the miocene, and that, consequently, the greenland ice sheet could not form prior to the late miocene uplift. the step (escarpment) between the ups and the lps follow the incision along the major valleys, whereas the pronounced topographical step across pdmf is caused by the differences in lithology. 11.3 fault offsets our data document active movements along the pdmf system during the early jurassic c1 and middle triassic c2 episodes. whereas denudation during these episodes reached far into the present-day basement terrains in the interior of the study area, the rift basins along the margin received the erosional products derived from the exhumation during these events. therefore, the pdmf acted as the separator between the uplifted areas to the west and the subsiding basins to the east of the fault system. our data document movements of the dombjerg fault during the mid-cretaceous c4 episode. the pdmf and adjacent fault systems were reactivated during the end-eocene c6 and the late miocene c8 episodes. afta data provide evidence of end-eocene block rotation of 600 m on hudson land, and there is a post-basalt throw along a major fault west of hold with hope of over 500 m. a pronounced increase in the estimated rock uplift since the late miocene across short distances over the pdmf indicates differential movements, allowing primarily miocene sediments to accumulate prior to their removal during late miocene uplift. the sediments thus accumulated in the coastal zone during a period of crustal extension prior to late miocene uplift during a compressional phase. the continuity of the lps on both sides of the pdmf in the south-western part of the study area indicates that this fault system was not active during the early pliocene c9 episode. 11.4 onshore-offshore correlation two important phases of progradation away from north-east greenland occurred after break-up. the first phase of progradation involving extensive mass wasting occurred after erosional incision along the shelf that happened in the interval between late eocene and mid-miocene. we suggest that the mass wasting offshore was due to uplift in the end-eocene c6 episode onshore, and thus that the horizon defining the erosional incision can be dated to about 35 ma. the next phase of massive shelf progradation occurred after the formation of the distinct imu. the imu marks the termination of synrift deposition in the deep-sea basins and the accelerated widening of the fram strait, and it is linked to mild compression and the onset of late miocene uplift onshore and massive shelf progradation. we therefore suggest that the uplift onshore and the subsequent progradation offshore began during the late miocene c8 episode, and thus that the imu can be dated to about 10 ma. 11.5 regionally synchronous episodes the timing of the regional episodes of uplift and erosion prior to break-up of the north-east atlantic defined in this study correlates closely with events in neighbouring regions. the three earliest episodes of uplift and denudation identified here also had a profound impact in north-west europe, and it is likely that the onset of these events was simultaneous in east greenland and southern scandinavia in the late carboniferous, the middle triassic and at the early–middle jurassic transition. exhumation most likely resulted from thermal uplift due to accumulation of heat beneath the supercontinent, pangaea, which preceded three phases of fragmentation of the supercontinent. 11.6 non-synchronous events episodes of uplift after break-up are not regionally synchronous across the atlantic. uplift and denudation http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 85 of 154 www.geusbul let in.org during the end-eocene c6 episode primarily affected arctic regions, including west and east greenland, with minor effects in southern scandinavia. early miocene uplift and erosion had a major impact in southern scandinavia, but there is no known counterpart in greenland. uplift and incision during the late miocene c8 episode initiated the formation of the modern relief in west and east greenland but has no known counterpart in southern scandinavia. coeval events to episodes c6 and c8 are reported from svalbard and the barents sea. early pliocene c9 uplift and incision affected both west and east greenland as well as southern scandinavia. we speculate that the forces driving (1) uplift during episode c6 in greenland are most likely linked with the stress that prevented further movements of the greenland plate relative to north america, (2) uplift during the early miocene in scandinavia probably originated from stress transmitted from orogenies on the eurasian plate, (3) uplift during episode c8 in greenland may be linked to changes in the absolute motion of north america and (4) uplift during episode c9 in greenland and scandinavia may be related to changes in upper-mantle flow related to the iceland plume. 11.7 episodic burial and exhumation as we have shown here, north-east greenland is a typical example of an epcm shaped by episodes of burial and exhumation both before and after break-up along the margin. in agreement with mckenzie’s theory of rifting, there is evidence for cretaceous and eocene postrift burial of the rift basins in north-east greenland. however, little remains of these deposits due to episodes of mid-cretaceous and cenozoic denudation that are not predicted by mckenzie’s theory. the events of uplift that followed the break-up of pangaea are, as yet, unexplained by conventional theories of basin development, as are the forces that lifted the peneplains that dominate the landscape of in north-east greenland to their present-day elevations. it is our hope that geoscientists with insight into the physical properties of the earth’s interior will now use our observations to explain vertical motions of the earth’s surface on timescales of tens of millions of years. acknowledgements many colleagues from the geological survey of greenland and denmark (geus) contributed with samples to this study; more than we can list here. christian knudsen collected many of the first samples analysed for this study during a sailboat expedition in 2008. jørgen bojesen-koefoed organised the expeditions to north-east greenland in 2008–2011, during which most of the samples were acquired, and he supported our work in many ways. kirsten hansen made apatite separates available for this study. niels henriksen (oscar) provided invaluable help in retrieving samples collected in past decades from geus’ core store. finn surlyk and sierd cloetingh provided insightful reviews of the manuscript. additional information funding statement the research was funded by a consortium of oil companies and geus. competing interests the authors declare no competing interests. author contributions pj: conceptualisation, data curation, investigation, methodology, project administration, writing – original draft, writing – review & editing. pfg: formal analysis, methodology, investigation, writing – original draft, writing – review & editing. jmb, mb, jrh: investigation, writing – review & editing. additional files five supplementary files are available at https://doi.org/10.22008/fk2/7tvnhb three appendices are available at the end of this monograph. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 86 of 154 www.geusbul let in.org ahnert, f. 1998: introduction to geomorphology, 352 pp. london: arnold. amante, c. & eakins, b.w. 2009: etopo1 1 arc-minute global relief 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(eds): petroleum geology of north westeurope. geological society, london, petroleum geology conference series 5. https://doi.org/10.1144/0050325 willerslev, e. et al. 2007: ancient biomolecules from deep ice cores reveal a forested southern greenland. science 317, 111–114. https://doi. org/10.1126/science.1141758 worsley, t.r., nance, d. & moody, j.b. 1984: global tectonics and eustasy for the past 2 billion years. marine geology 58, 373–400. https://doi. org/10.1016/0025-3227(84)90209-3 zachos, j.c., pagani, m., sloan, l., thomas, e. & billups, k. 2001: trends, rhythms, and aberrations in global climate 65 ma to present. science 292, 686–693. https://doi.org/10.1126/science.1059412 ziegler, p.a. 1990: geological atlas of western and central europe, 239 pp. the hague: shell international petroleum matschappij b.v. https:// doi.org/10.2307/3060311 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 93 of 154 www.geusbul let in.org techniques for extracting thermal history constraints from afta data were discussed in detail by green & duddy (2012) and green et al. (2013). a brief outline based on application to sedimentary rocks is provided here. similar principles apply to basement samples which have been buried by sedimentary cover. because afta data are dominated by maximum postdepositional temperatures, the data preserve no evidence of the history between deposition and the onset of cooling from the maximum post-depositional temperature. as such, the history between successive cooling episodes separated by periods of heating cannot be resolved. therefore, it is not possible to constrain the entire history from the beginning of track retention. instead, we focus on defining the palaeothermal maximum and subsequent palaeothermal peaks, within an overall framework of episodic heating and cooling, assuming heating and cooling rates of 1°c/myr and 10°c/myr, respectively. an order of magnitude change in heating rate is equivalent to around a 10°c difference in the palaeotemperature required to produce a given degree of annealing. we propose that thermal histories involving episodic heating and cooling are more geologically realistic than slow, monotonic cooling, even in basement rocks (green & duddy 2012; green et al. 2013, 2018). this is based on a combination of afta data and geological evidence where afta defines episodes of cooling/exhumation while remnants of former sedimentary cover define times when underlying rocks were at the surface. such situations provide indisputable evidence of repeated burial and exhumation. green et al. (2013, 2018) argued that such histories are more generally applicable in a wider variety of settings. thermal history solutions are extracted from data in each sample by comparing measured data with values predicted from candidate thermal histories, involving up to three episodes of heating and cooling (three being the maximum number of cooling episodes that can be obtained from afta data in most cases). by varying the maximum palaeotemperature and the onset of cooling in each episode (using assumed heating and cooling rates of 1°c/myr and 10°c/myr, respectively), we defined the range of conditions, which predict apatite fission-track ages, track length distributions and their variation with wt% cl consistent with the measured data to within 95% confidence limits. the approach is based on likelihood theory and principles similar to those outlined by gallagher (1995). we emphasise that unlike other approaches, we do not attempt to constrain the entire history since the onset of track retention, because much of the history is not recorded in the data (green & duddy 2012; green et al. 2013). instead, results are presented in terms of intervals representing 95% confidence limits on the maximum (peak) palaeotemperature and the time at which cooling from that palaeotemperature began in up to three intervals (appendix 1.1). it is important to appendix 1 thermal history constraints from afta data in individual samples and definition of regional palaeothermal episodes 100 80 60 40 20 0 120 250 200 50150 100 0 t2: 95% c.i. on maximum temperature t3 t3 t1 t1 t2: 95% c.i. on onset of cooling te m pe ra tu re (° c ) time (ma) fig. 9 deposition appendix 1.1 illustration of thermal history solutions extracted from afta. results are presented in terms of up to three palaeothermal episodes, that is, when a rock sample was hotter than it is today. in the case shown, the rock sample was exposed at the surface during deposition of sediments as indicated by the small box. note that the palaeotemperature constraint t1 is a minimum estimate since this is when apatites begin to retain tracks. in most situations, up to three episodes can be defined from afta, due to various factors including the natural spread in track lengths of a single population of tracks and the rapid decrease in the rate of annealing from c. 110°c to below 60°c. t1, t2, t3: time intervals during which cooling from the peak palaeotemperature began. t1, t2, t3: palaeotemperature intervals. c.i.: confidence limits. colours are used to illustrate attribution to regional episodes, although those used here are purely schematic. reproduced from japsen et al. (2021, this volume). thermal history interpretation of afta data: principles http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 94 of 154 www.geusbul let in.org aconstrain the thermal histories derived from afta data by geological evidence. for outcrops of sedimentary rocks, the most fundamental constraint is the depositional age, which defines a time when the sample was at surface temperature. thermal history solutions are derived within a framework fixed at the depositional age and the present-day. for basement samples, the most reliable solutions are obtained when cover rocks are present locally, as these define a time of surface exposure. where cover rocks are not present, a crystallisation or metamorphic age can be used. in such cases the solutions are less well constrained. the kinetics of fission track annealing in apatite vary significantly with chlorine content (wt% cl). acceptable thermal history solutions must predict the measured apatite fission-track age length distribution and the variation of these with wt% cl in each sample. wt% cl is measured in each apatite analysed, using an electron microprobe (see appendix a in the geotrack reports). where three palaeothermal episodes (i.e. periods when samples cooled from elevated temperatures) are defined in a single sample, the earliest episode is usually defined primarily from the fission-track age data and represents the episode in which samples cooled below c. 110°c and began to retain tracks. the two more recent episodes are defined primarily from the distribution of track lengths. grains with higher wt% cl are more retentive, so samples containing a spread of wt% cl may also define earlier cooling events. results from north-east greenland sample details and afta data are listed in appendices 1.2 and 1.3, respectively. the thermal history constraints extracted from the afta data in each sample are summarised in appendix 1.4. vr data and corresponding maximum palaeotemperatures are listed in appendix 1.5. the stratigraphic ages in these appendices are assigned following gradstein et al. (2012), and the tables are provided online in supplementary file s4. appendix 1.6 shows a comparison of the constraints (95% confidence intervals) on the onset of cooling in up to three palaeothermal episodes in each sample based on the thermal-history solutions extracted from the afta data in each sample. based on the general uniformity of the data, we have sought to define the minimum number of regionally synchronous cooling episodes, which can explain the results in all these samples. this synthesis has been undertaken within the context of the locations of individual samples, such that similar magnitude events in adjacent samples are interpreted in terms of a common episode. on this basis, the timing constraints defined from afta data, listed in appendix 1.4, are interpreted as defining ten episodes of regional cooling (c0–c9) dating from the late palaeozoic to the early pliocene (table 1). note that the time intervals in the table define when cooling began in each episode, and we do not imply that all cooling took place solely within these intervals. it should also be emphasised that because the afta data only define the palaeothermal peak and not the history prior to the onset of cooling, the total amount of cooling in each episode and the degree of re-heating leading to the subsequent palaeothermal peak is not controlled by the data. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 95 of 154 www.geusbul let in.org appendix 1.2 sample details. also available online in supplementary file s4 sample number stratigraphic division stratigraphic age (ma) rock type latitude longitude elevation (m a.s.l.) location gc103-1 permian, l 255 sandstone 71.676 –22.908 200 jameson land n gc103-2 permian, l 255 sandstone 71.657 –22.957 200 jameson land n gc103-3 jurassic, l 155 sandstone 70.618 –22.787 660 jameson land s gc103-4 jurassic, e–m 190 sandstone 70.863 –22.716 400 jameson land s gc103-6 permian, l 255 sandstone 71.800 –24.102 1050 schuchert dal gc159-1 permian, l 255 sandstone 71.648 –24.662 550 schuchert dal gc159-2 permian, l 255 sandstone 71.614 –24.543 260 schuchert dal gc159-3 triassic, e 251–247 sandstone 71.583 –23.008 250 jameson land n gc159-4 permian, l 260–250 sandstone 71.714 –22.702 400 jameson land n gc202-1 permian, l 260–250 arkose 71.534 –24.715 820 schuchert dal gc202-2 permo–carboniferous 360–245 71.753 –24.338 700 schuchert dal gc202-3 permian, l 255 sandstone 71.785 –24.427 820 schuchert dal gc202-4 permian, l 260–250 sandstone 71.763 –24.381 1000 schuchert dal gc202-5 triassic, m 245–240 sandstone 71.698 –24.195 1000 schuchert dal gc202-6 triassic, m 245–240 sandstone 71.719 –24.266 1310 schuchert dal gc202-7 triassic, l 235–208 sandstone 71.722 –24.131 1000 schuchert dal gc202-9 jurassic, e 208–178 sandstone 71.589 –23.731 900 jameson land s gc202-11 triassic, l 235–208 mudstone/ sandstone 70.997 –22.575 150 jameson land s gc202-12 triassic, l 235–208 mudstone/ sandstone 71.160 –22.503 130 jameson land s gc202-16 permian, l 255 mudstone/ sandstone 71.737 –24.307 720 schuchert dal gc522-1 triassic 245–208 sediment 72.567 –22.417 500 traill ø gc522-2 devonian 416–359 sediment 72.583 –24.250 30 traill ø gc522-3 devonian 416–359 sediment 72.400 –25.833 0 traill ø gc522-4 triassic 245–208 sediment 72.283 –23.000 140 traill ø gc522-5 carboniferous 359–299 sediment 72.200 –24.000 70 traill ø gc522-6 carboniferous 359–299 sediment 72.583 –24.250 150 traill ø gc522-7 jurassic, l 161–146 sediment 74.450 –20.252 150 wollaston forland gc522-8 cretaceous. e 146–132 sediment 74.585 –20.475 300 wollaston forland gc522-9 l. carboniferous 313–303 sediment 72.443 –23.801 140 traill ø gc522-10 bathonian 166–161 sediment 72.366 –23.321 230 traill ø gc522-11 callovian–bathonian 166–157 sediment 72.470 –22.411 300 traill ø gc1016-3 cretaceous, e 146–100 sandstone 74.519 –20.451 653.5 wollaston forland gc1016-4 silurian 430 monzonite 71.054 –26.064 5 inland s gc1016-5 devonian 416–359 conglomerate 70.968 –28.021 100 inland s gc1016-6 precambr./caled. >400 migmatite 70.420 –27.806 1 milne land gc1016-7 precambr./caled. >400 gneiss 70.450 –26.251 2 milne land gc1016-8 precambr./caled. >400 sandstone 72.874 –25.127 120 ella ø gc1016-9 devonian 416–359 conglomerate 72.789 –25.036 1 ella ø gc1016-10 carboniferous 359–299 sandstone 72.270 –23.927 1 jameson land n gc1016-11 permian 299–251 sandstone 72.144 –23.693 50 jameson land n gc1016-13 precambr./caled. >400 sandstone 72.444 –25.097 1 inland c gc1016-14 precambr./caled. >400 sandstone 72.460 –24.927 1 inland c gc1016-15 precambr./caled. >400 quartzite 73.330 –25.287 1 inland c gc1016-16 devonian 416–359 sandstone 73.348 –23.735 1 ymer ø gc1016-17 devonian 416–359 sandstone 73.309 –23.504 1 ymer ø http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 96 of 154 www.geusbul let in.org appendix 1.2 sample details. also available online in supplementary file s4 (continued) sample number stratigraphic division stratigraphic age (ma) rock type latitude longitude elevation (m a.s.l.) location gc1016-18 devonian 416–359 sandstone 73.258 -23.280 1 ymer ø gc1016-19 precambr./caled. >400 gneiss 76.774 –18.717 1 germania land gc1016-20 precambr./caled. >400 gneiss 74.410 –20.921 1 clavering ø gc1016-21 eocene 49–47 sandstone 69.447 –24.148 303 kap dalton gc1016-22 eocene 49–47 sandstone 69.481 –24.141 260 kap dalton gc1016-23 paleocene 65–56 sandstone 74.417 –19.200 64 wollaston forland gc1016-24 carboniferous 359–299 sandstone 73.923 –22.275 756 hudson land gc1016-25 carboniferous? 359–299 sandstone 73.947 –22.165 189 hudson land gc1016-26 carboniferous 359–299 sandstone 73.933 –22.244 479 hudson land gc1016-27 jurassic–cretaceous 200–65 sandstone 70.502 –23.062 250 jameson land s gc1016-28 triassic 251–200 sandstone 70.499 –23.049 95 jameson land s gc1016-29 jurassic–cretaceous 200–65 sandstone 72.454 –23.299 730 traill ø gc1016-30 jurassic–cretaceous 200–65 sandstone 72.446 –23.298 1075 traill ø gc1016-32 jurassic, e 200–176 sandstone 72.460 –23.235 323 traill ø gc1016-33 jurassic, e 200–176 sandstone 72.460 –23.278 460 traill ø gc1016-35 cretaceous, e 146–100 sandstone 72.952 –23.081 300 geographical society ø gc1016-36 cretaceous, e 146–100 sandstone 72.934 –23.057 370 geographical society ø gc1016-37 triassic 251–200 sandstone 72.933 –23.013 850 geographical society ø gc1016-38 triassic 251–200 sandstone 73.928 –21.174 381 hold with hope gc1016-39 triassic 251–200 sandstone 73.920 –21.160 751 hold with hope gc1016-41 cretaceous, l 112–65 sandstone 73.707 –20.543 21 hold with hope gc1016-42 ordovician–devonian 488–359 gneiss 70.605 –26.000 41 milne land gc1016-43 ordovician 488–444 granite 70.627 –26.082 362 milne land gc1016-44 jurassic 200–146 sandstone 70.616 –26.083 759 milne land gc1016-45 ordovician 488–444 gneiss 70.598 –26.043 458 milne land gc1016-46 precambr./caled. >400 gneiss 70.916 –26.250 1858 milne land gc1016-47 precambr./caled. >400 gneiss 70.936 –26.278 2 milne land gc1016-48 precambr./caled. >400 gneiss 74.389 –20.240 28 wollaston forland gc1016-49 carboniferous 359–299 sand/ conglomerate 74.245 –21.875 846 clavering ø gc1016-50 carboniferous 359–299 sand/ conglomerate 74.244 –21.880 755 clavering ø gc1016-51 carboniferous 359–299 sand/ conglomerate 74.242 –21.887 574 clavering ø gc1016-52 carboniferous 359–299 sand/ conglomerate 74.230 –21.870 499 clavering ø gc1016-53 carboniferous 359–299 sand/ conglomerate 74.242 –21.894 501 clavering ø gc1016-54 carboniferous 359–299 sand/ conglomerate 74.238 –21.923 248 clavering ø gc1016-55 carboniferous 359–299 sand/ conglomerate 74.234 –21.964 52 clavering ø gc1016-57 precambr./caled. >400 gneiss 73.828 –21.948 1208 hudson land gc1016-58 precambr./caled. >400 gneiss 73.826 –21.909 958 hudson land gc1016-59 precambr./caled. >400 gneiss 73.817 –21.934 720 hudson land gc1016-60 precambr./caled. >400 gneiss 73.811 –21.919 510 hudson land gc1016-61 precambr./caled. >400 gneiss 73.799 –21.918 305 hudson land gc1016-62 precambr./caled. >400 gneiss 73.827 –21.850 123 hudson land http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 97 of 154 www.geusbul let in.org appendix 1.2 sample details. also available online in supplementary file s4 (continued) sample number stratigraphic division stratigraphic age (ma) rock type latitude longitude elevation (m a.s.l.) location gc1016-63 precambr./caled. >400 gneiss 73.743 –22.101 569 hudson land gc1016-65 precambr./caled. >400 gneiss 73.712 –22.112 210 hudson land gc1016-66 permian 299–251 sandstone 74.046 –21.787 0 hold with hope gc1016-67 triassic 251–200 sandstone 74.028 –21.781 292 hold with hope gc1016-68 triassic 251–200 sandstone 74.005 –21.802 323 hold with hope gc1016-69 triassic 251–200 sandstone 73.998 –21.880 94 hold with hope gc1016-70 carboniferous 359–299 sandstone 73.918 –22.259 959 hudson land gc1016-72 precambr./caled. >400 basement 71.354 –21.906 445 liverpool land gc1016-73 precambr./caled. >400 granite 71.416 –22.154 550 liverpool land gc1077-1 precambr./caled. >400 basement 76.283 –25.900 1800 inland n gc1077-3 precambr./caled. >400 basement 76.783 –25.250 900 inland n gc1077-4 precambr./caled. >400 basement 75.233 –21.783 100 hochstetter forland w gc1077-5 precambr./caled. >400 basement 76.000 –22.100 340 dove bugt gc1077-6 precambr./caled. >400 basement 75.983 –20.400 0 dove bugt gc1077-8 precambr./caled. >400 basement 76.333 –20.600 775 dove bugt gc1077-9 precambr./caled. >400 basement 76.917 –21.417 20 dove bugt gc1077-10 precambr./caled. >400 basement 76.967 –20.483 200 dove bugt gc1077-11 precambr./caled. >400 basement 77.533 –21.017 510 germania land gc1077-12 barremian 130–125 sandstone 76.380 –18.728 54 store koldewey gc1077-13 barremian 130–125 sand, unconsolidated 76.427 –18.705 34 store koldewey gc1077-15 jurasic, m 176–161 sandstone 76.118 –18.532 108 store koldewey gc1077-16 devonian, e 416–398 gneiss 76.117 –18.533 105 store koldewey gc1077-17 callovian 165–161 sandstone 75.190 –19.998 2 hochstetter forland gc1077-18 neoproterozoic 800 basement 72.229 –25.333 1785 stauning alper gc1077-20 neoproterozoic 800 basement 72.257 –25.359 373 stauning alper gc1077-21 neoproterozoic 800 basement 72.267 –25.361 12 stauning alper gc1077-22 devonian, e 416–398 gneiss, metasediment 74.880 –20.202 765 kuhn ø gc1077-23 devonian, e 416–398 granite 74.880 –20.214 572 kuhn ø gc1077-24 devonian, e 416–398 granite 74.873 –20.227 400 kuhn ø gc1077-25 devonian, e 416–398 granite 74.883 –20.237 225 kuhn ø gc1077-26 devonian, e 416–398 granite 74.885 –20.246 100 kuhn ø gc1077-27 devonian, e 416–398 basement 74.290 –22.477 1261 payer land gc1077-28 devonian, e 416–398 basement 74.304 –22.514 939 payer land gc1077-29 devonian, e 416–398 basement 74.316 –22.560 640 payer land gc1077-30 devonian, e 416–398 basement 74.316 –22.589 479 payer land gc1077-31 devonian, e 416–398 basement 74.323 –22.577 171 payer land gc1077-32 devonian, e 416–398 basement 74.313 –22.458 32 payer land gc1077-33 devonian, e 416–398 basement 70.837 –22.017 1277 liverpool land gc1077-34 devonian, e 416–398 basement 70.834 –22.028 1058 liverpool land gc1077-35 devonian, e 416–398 basement 70.817 –22.052 600 liverpool land gc1077-36 devonian, e 416–398 basement 70.844 –21.895 317 liverpool land gc1077-37 callovian–oxfordian 165–156 sandstone 71.463 –23.328 1217 jameson land s gc1077-38 bajocian–bathonian 172–165 sandstone 71.391 –23.339 857 jameson land s gc1077-39 pliensbachian–toarcian 190–176 fine-grained sandstone 71.481 –23.372 469 jameson land s http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 98 of 154 www.geusbul let in.org appendix 1.2 sample details. also available online in supplementary file s4 (continued) sample number stratigraphic division stratigraphic age (ma) rock type latitude longitude elevation (m a.s.l.) location gc1077-40 paleocene–eocene 66–34 unconsolidated sand 74.028 –21,668 739 hold with hope gc1077-41 midproterozoic 540 granite 71.418 –25.341 0 stauning alper gc1077-42 silurian 444–416 granite 73.270 –22.199 160 gauss halvø gc1077-43 precambr./caled. >400 basement 73.543 –21.954 87 gauss halvø gc1077-44 precambr./caled. >400 basement 73.540 –21.998 409 gauss halvø gc1077-45 permian 299–251 sandstone 73.541 –22.017 455 gauss halvø gc1077-46 triassic 251–200 sandstone 73.535 –22.064 621 gauss halvø gc1077-47 carboniferous 359–299 conglomerate 74.413 –21.867 120 payer land gc1077-49 precambr./caled. >400 basement 74.471 –21.943 413 payer land gc1077-50 carboniferous 359–299 sandstone 73.147 –23.084 255 ymer ø gc1077-51 devonian 416–359 sandstone 73.174 –23.182 103 ymer ø gc1077-52 tertiary 36 syenite 72.202 –22.730 1018 traill ø gc1077-54 tertiary 36 syenite 72.190 –22.713 689 traill ø gc1077-55 tertiary 36 syenite 72.184 –22.697 128 traill ø gc1077-56 midproterozoic >400 paragneiss 71.587 –27.059 0 inland s gc1077-57 archaeic >400 gneiss 71.621 –27.602 0 inland s gc1077-58 precambr./caled. >400 gneiss 71.161 –28.708 800 inland s gc1077-59 precambr./caled. >400 schist 70.464 –28.322 500 gåseland gc1077-60 proterozoic, m–l 540 gneiss 75.082 –21.320 0 hochstetter forland w gc1077-61 palaeoproterozoic >540 gabbro 75.032 –21.666 360 hochstetter forland w gc1077-62 precambr./caled. >400 semipelite 73.742 –29.325 2165 inland c gc1077-63 palaeoproterozoic >540 granite 72.314 –26.858 800 inland c gc1077-64 palaeoproterozoic >540 gneiss 72.096 –28.836 930 inland c gc1077-65 neoproterozoic >540 granite 73.483 –26.504 870 inland c gc1077-66 neoproterozoic >540 granite 73.851 –26.152 650 inland c gc1077-67 paleoprotorozoic? 540 gneiss 73.468 –27.676 280 inland c gc1077-68 precambr./caled. >400 metaquarts diorite 73.194 –27.660 500 inland c gc1077-69 silurian 444–416 granite 73.062 –28.116 2110 inland c gc1077-70 proterozoic, m–l 540 gneiss 74.729 –22.626 1150 payer land gc1077-71 neoproterozoic >540 gneiss 74.253 –25.242 950 inland c gc1077-72 paleocene–eocene, e 66–50 sandstone 74.565 –18.984 402 sabine ø gc1077-73 cretaceous, e 146–100 conglomerate 74.584 –19.339 150 wollaston forland gc1077-74 devonian, e 416–398 granite 74.584 –19.334 154 wollaston forland gc1077-75 cretaceous 146–66 sandstone 74.596 –19.542 340 wollaston forland gc1077-76 triassic, l 228–200 sandstone 70.719 –22.661 110 jameson land s gc1077-77 jurassic, l 161–146 sandstone 70.658 –22.699 717 jameson land s gc1077-78 rhaetian–sinemurian 204–189 sandstone 70.745 –22.674 40 jameson land s gc1077-79 triassic 251–200 sandstone 73.933 –21.215 150 hold with hope gc1077-80 cretaceous? 145–65 sandstone 74.359 –20.576 457 clavering ø gc1077-81 cretaceous 146–66 sandstone 74.378 –20.593 303.5 clavering ø gc1077-82 devonian, e 416–398 gneiss 74.378 –20.593 303 clavering ø gc1077-83 devonian, e 416–398 basement 76.232 –18.585 107 store koldewey gc1077-84 cretaceous, e 146–100 sandstone 74.346 –19.291 58 wollaston forland gc1077-85 eocene, e 55 sandstone 73.,755 –21.940 320 hudson land gc1077-86 precambr./caled. >400 basement 74.169 –20.771 778 clavering ø http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 99 of 154 www.geusbul let in.org appendix 1.2 sample details. also available online in supplementary file s4 (continued) sample number stratigraphic division stratigraphic age (ma) rock type latitude longitude elevation (m a.s.l.) location gc1077-87 triassic 251–200 sandstone 74.169 –20.771 778 clavering ø gc1077-88 precambr./caled. >400 basement 74.342 –20.086 327 wollaston forland gc1077-89 cretaceous, e 146–100 conglomerate 74.582 –20.751 500 dombjerg gc1077-90 precambr./caled. >400 basement 74.582 –20.751 500 dombjerg gc1077-91 precambr./caled. >400 basement 74.546 –20.789 1446 dombjerg gc1077-92 precambr./caled. >400 basement 74.534 –20.683 234.696 dombjerg gc1077-93 precambr./caled. >400 basement 74.335 –20.793 820 clavering ø gc1077-94 precambr./caled. >400 basement 74.336 –20.783 890 clavering ø gc1077-95 precambr./caled. >400 basement 74.350 –20.745 1233 clavering ø gc1077-96 paleocene, sub-basalt 65–56 sandstone 74.350 –20.745 1233 clavering ø gc1077-97 precambr./caled. >400 basement 74.342 –20.701 427 clavering ø gc1077-98 jurassic, l 161–146 sandstone 74.645 –20.054 50 wollaston forland gc1077-99 devonian, e 416–398 basement 74.645 –20.059 100 wollaston forland gc1077-100 miocene, e 23–16 sill 74.520 –18.751 81 hvalros ø gc1077-101 valanginian? 140–136 sandstone 76.297 –18.634 77 store koldewey gc1077-102 devonian, e 416–398 gneiss 76.297 –18.634 77 store koldewey gc1077-103 devonian, e 416–398 gneiss 76.169 –18.659 663 store koldewey gc1077-104 devonian, e 416–398 gneiss 76.318 –18.777 524 store koldewey gc1077-105 paleocene 65–56 sandstone 74.344 –19.360 420 wollaston forland gc1104-1 precambr./caled. >400 metagabbro 77.138 –18.573 280 germania land gc1104-2 precambr./caled. >400 gneiss 76.697 –19.844 192.5 dove bugt gc1104-3 precambr./caled. >400 orthogneiss 76.288 –20.352 830 dove bugt gc1104-4 precambr./caled. >400 granite 70.938 –25.580 600 milne land gc1104-5 precambr./caled. >400 granite 70.842 –26.144 1270 milne land gc1104-6 precambr./caled. >400 migmatite neosome 70.855 –25.855 1280 milne land gc1104-7 precambr./caled. >400 granite 71,480 –25.228 1090 stauning alper gc1104-8 precambr./caled. >400 granite 71.468 –25.304 300 stauning alper gc1104-9 precambr./caled. >400 granite 71.468 –25.304 1450 stauning alper gc1104-10 precambr./caled. >400 migmatite 71.960 –24.781 1280 stauning alper gc1104-12 precambr./caled. >400 gneiss 72.003 –25.528 20 stauning alper gc1104-13 precambr./caled. >400 granite 72.132 –25.519 50 stauning alper gc1104-14 precambr./caled. >400 granite 73.606 –25.846 1100 inland c gc1104-15 campanian? 83.5–70.6 sandstone 72.849 –22.441 656 geographical society ø gc1104-16 triassic 251–200 sandstone 72.727 –23.144 570 traill ø gc1104-17 jurassic? 200–146 conglomerate 72.727 –23.144 570 traill ø gc1104-18 jurassic? 199.6–145.5 sandstone 72.459 –22.384 471 traill ø gc1104-19 triassic 251–199 sandstone 72.751 –23.306 1069 traill ø gc1104-20 precambr./caled. >400 basement 76.105 –18.502 48 store koldewey gc1104-21 precambr./caled. >400 porphyritic band 70.902 –25.376 2 milne land gc1104-22 precambr./caled. >400 basement 76.334 –18.667 54 store koldewey gc1104-23 valanginian 140.2–136.4 sandstone 76.334 –18.667 54 store koldewey gc1104-24 santonian 85.8–83.5 sandstone 73.707 –20.546 –5 hold with hope, nanok bh for definition of sub-regions inland n, inland c and inland s, see figure 11a. caled.: caledonian; e: early; l: late; m: middle; precambr.: precambrian. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 100 of 154 www.geusbul let in.org appendix 1.3 afta data. also available online in supplementary file s4 sample number stratigraphic age a (ma) ρd b (106 tracks /cm2) ρs b (106 tracks /cm2) ρi b (106 tracks /cm2) p(χ2) c (number of grains) fission-track age d (ma) mean track length e (μm) (number of lengths) standard deviation f (μm) gc103-1 255 1.486 (2357) 0.370 (236) 6.359 (4054) 29 (20) 17.0 ± 1.2 12.6 ± 0.2 (102) 1.98 gc103-2 255 1.488 (2357) 0.279 (145) 5.464 (2840) 49 (20) 14.9 ± 1.3 12.5 ± 0.2 (107) 2.28 gc103-3 155 1.490 (2357) 1.106 (456) 2.737 (1128) <1 (20) 119.5 ± 17.7* 10.3 ± 0.2 (106) 2.36 gc103-4 190 1.492 (2357) 0.260 (234) 1.076 (970) 52 (20) 70.3 ± 5.5 11.1 ± 0.2 (100) 2.49 gc103-6 255 1.496 (2357) 1.488 (618) 3.814 (1584) <1 (20) 153.4 ± 17.5* 10.2 ± 0.3 (102) 2.62 gc159-1 255 1.391 (2205) 0.515 (302) 3.364 (1971) 79 (20) 40.4 ± 2.7 12.7 ± 0.2 (100) 1.58 gc159-2 255 1.392 (2205) 0.510 (438) 5.255 (4514) <1 (20) 27.0 ± 2.5* 12.4 ± 0.2 (75) 1.97 gc159-3 251–247 1.394 (2205) 0.139 (120) 2.074 (1795) <1 (20) 17.7 ± 1.7 13.9 ± 0.3 (14) 1.05 gc159-4 260–250 1.396 (2205) 0.371 (327) 5.371 (4732) <1 (20) 17.7 ± 1.9* 13.3 ± 0.2 (61) 1.44 gc202-1 260–250 1.440 (2198) 1.617 (806) 3.363 (1676) <1 (20) 132.8 ± 13.0* 11.3 ± 0.3 (106) 2.62 gc202-2 360–245 1.433 (2198) 0.386 (223) 2.702 (1559) 57 (20) 40.1 ± 3.1 11.9 ± 0.2 (101) 2.15 gc202-3 255 1.426 (2198) 0.708 (366) 3.912 (2021) <1 (20) 54.7 ± 8.6* 11.0 ± 0.2 (101) 2.23 gc202-4 260–250 1.419 (2198) 0.690 (420) 2.746 (1671) <1 (20) 65.6 ± 10.0* 11.4 ± 0.3 (103) 2.58 gc202-5 245–240 1.412 (2198) 0.847 (464) 2.169 (1189) <1 (20) 98.6 ± 12.3* 10.9 ± 0.2 (103) 2.27 gc202-6 245–240 1.404 (2198) 1.592 (627) 2.498 (984) <1 (20) 120.4 ± 19.4* 10.8 ± 0.2 (104) 2.02 gc202-7 235–208 1.390 (2198) 1.115 (228) 7.525 (1539) 60 (11) 40.5 ± 3.3* 12.7 ± 0.2 (51) 1.59 gc202-9 208–178 1.383 (2198) 0.661 (442) 2.608 (1743) <1 (20) 66.6 ± 6.3* 12.1 ± 0.2 (103) 2.22 gc202-11 235–208 1.375 (2198) 0.401 (309) 2.544 (1958) <1 (20) 44.7 ± 4.4* 12.4 ± 0.2 (113) 2.21 gc202-12 235–208 1.368 (2198) 0.144 (113) 1.265 (991) 3 (20) 29.6 ± 4.8* 12.6 ± 0.3 (51) 2.25 gc202-16 255 1.361 (2198) 0.965 (531) 8.544 (4699) <1 (20) 30.3 ± 2.3* 12.2 ± 0.2 (123) 2.42 gc522-1 245–208 1.656 (2712) 0.286 (170) 1.766 (1049) <1 (20) 37.3 ± 9.7* 13.5 ± 0.3 (44) 1.92 gc522-2 416–359 1.666 (2712) 0.830 (267) 2.220 (714) <1 (20) 104.4 ± 14.7* 11.4 ± 0.3 (55) 2.36 gc522-3 416–359 1.676 (2712) 2.220 (1031) 2.799 (1300) 3 (20) 229.5 ± 14.4* 12.3 ± 0.2 (105) 2.01 gc522-4 245–208 1.685 (2712) 0.112 (65) 1.902 (1105) 41 (19) 17.6 ± 2.3 10.7 ± 1.3 (11) 4,21 gc522-5 359–299 1.467 (2300) 0.151 (101) 3.090 (2061) <1 (20) 11.3 ± 2.1* 12.4 ± 0.3 (58) 2.44 gc522-6 359–299 1.466 (2300) 1.519 (347) 3.607 (824) <1 (19) 111.3 ± 14.7* 10.6 ± 0.3 (78) 2.38 gc522-7 161–146 1.464 (2300) 2.006 (654) 3.006 (980) 16 (20) 173.6 ± 10.0 10.4 ± 0.3 (103) 3.06 gc522-8 146–132 1.462 (2300) 1.975 (1065) 2.129 (1148) 81 (20) 239.8 ± 12.2 11.2 ± 0.2 (105) 1.89 gc522-9 313–303 1.267 (2088) 0.476 (191) 5.562 (2233) <1 (20) 20.6 ± 4.4* 12.1 ± 0.4 (45) 2.87 gc522-10 166–161 1.257 (2088) 0.183 (79) 2.591 (1117) 10 (20) 16.1 ± 2.7* 11.1 ± 0.4 (59) 2.79 gc522-11 166–157 1.247 (2088) 0.308 (92) 2.476 (740) <1 (20) 24.1 ± 8.5* 10.4 ± 0.3 (103) 3.11 gc1016-3 146–100 1.327 (2127) 3.045 (1483) 4.609 (2245) <1 (20) 158.2 ± 13.9* 12.9 ± 0.1 (101) 1.21 gc1016-4 430 1.331 (2127) 2.517 (1902) 3.664 (2769) <1 (20) 168.4 ± 10.4* 13.4 ± 0.1 (103) 1.35 gc1016-5 416–359 1.334 (2127) 2.691 (1629) 3.082 (1866) <1 (20) 218.8 ± 17.3* 12.4 ± 0.1 (102) 1.33 gc1016-6 >400 1.338 (2127) 1.649 (553) 2.045 (686) 46 (20) 201.9 ± 12.7 13.1 ± 0.2 (49) 1.37 gc1016-7 >400 1.341 (2127) 2.207 (1942) 3.078 (2708) <1 (20) 179.2 ± 12.6* 12.6 ± 0.1 (101) 1.25 gc1016-8 >400 1.345 (2127) 0.565 (95) 1.053 (177) 20 (11) 135.9 ± 17.6 13.7 ± 0.4 (6) 0.89 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 101 of 154 www.geusbul let in.org appendix 1.3 afta data. also available online in supplementary file s4 (continued) sample number stratigraphic age a (ma) ρd b (106 tracks /cm2) ρs b (106 tracks /cm2) ρi b (106 tracks /cm2) p(χ2) c (number of grains) fission-track age d (ma) mean track length e (μm) (number of lengths) standard deviation f (μm) gc1016-9 416–359 1.348 (2127) 1.533 (551) 2.346 (843) <1 (20) 164.3 ± 14.4* 13.3 ± 0.2 (55) 1.28 gc1016-10 359–299 1.352 (2127) 0.232 (248) 3.496 (3738) 4 (20) 16.7 ± 1.6* 13.5 ± 0.3 (46) 1.70 gc1016-11 299–251 1.356 (2127) 0.218 (120) 3.483 (1920) <1 (20) 19.4 ± 6.9* 11.6 ± 0.7 (9) 2.22 gc1016-13 >400 1.359 (2127) 1.152 (530) 2.200 (1012) <1 (20) 129.8 ± 14.5* 10.1 ± 0.3 (26) 1.76 gc1016-14 >400 1.363 (2127) 1.033 (303) 1.933 (567) <1 (20) 138.9 ± 17.2* 11.6 ± 0.5 (14) 2.02 gc1016-15 >400 1.366 (2127) 1.041 (323) 1.724 (535) <1 (15) 156.8 ± 18.2* 12.5 ± 0.4 (19) 1.72 gc1016-16 416–359 1.370 (2127) 0.969 (643) 1.776 (1179) <1 (20) 131.7 ± 14.8* 11.6 ± 0.2 (80) 1.59 gc1016-17 416–359 1.373 (2127) 1.812 (1615) 4.507 (4016) <1 (20) 104.2 ± 8.2* 11.3 ± 0.2 (79) 1.80 gc1016-18 416–359 1.284 (1945) 0.431 (225) 1.861 (971) <1 (20) 58.2 ± 8.0* 11.1 ± 0.2 (108) 2.32 gc1016-19 >400 1.277 (1945) 0.782 (281) 1.058 (380) 100 (20) 182.9 ± 15.4 12.8 ± 0.1 (107) 1.46 gc1016-20 >400 1.270 (1945) 0.993 (590) 1.397 (830) 48 (20) 175.0 ± 10.7 12.7 ± 0.1 (112) 1.58 gc1016-21 49–47 1.264 (1945) 1.496 (913) 2.343 (1430) <1 (20) 144.4 ± 16.6* 11.8 ± 0.1 (133) 1.66 gc1016-22 49–47 1.257 (1945) 0.092 (105) 0.425 (485) 24 (20) 53.2 ± 5.9 13.3 ± 0.3 (81) 2.59 gc1016-23 65–56 1.250 (1945) 1.824 (900) 2.345 (1157) <1 (20) 173.3 ± 15.0* 11.5 ± 0.2 (105) 1.85 gc1016-24 359–299 1.243 (1945) 1.792 (662) 2.802 (1035) 22 (13) 157.6 ± 11.1* 12.5 ± 0.2 (113) 1.87 gc1016-25 359–299 1.236 (1945) 1.429 (706) 2.522 (1246) <1 (20) 136.3 ± 12.2* 11.9 ± 0.2 (122) 1.74 gc1016-26 359–299 1.229 (1945) 1.930 (1007) 2.682 (1399) 2 (20) 171.6 ± 8.7 12.1 ± 0.2 (106) 1.68 gc1016-27 200–65 1.223 (1945) 2.075 (658) 3.140 (996) <1 (20) 148.8 ± 15.9* 10.0 ± 0.2 (113) 2.00 gc1016-28 251–200 1.216 (1945) 0.599 (217) 1.995 (723) <1 (20) 73.5 ± 13.6* 10.1 ± 0.2 (101) 2.37 gc1016-29 200–65 1.209 (1945) 0.476 (313) 2.340 (1540) <1 (20) 34.7 ± 7.9* 11.9 ± 0.3 (68) 2.24 gc1016-30 200–65 1.202 (1945) 0.390 (279) 2.698 (1932) <1 (20) 34.1 ± 3.7* 12.3 ± 0.2 (108) 2.38 gc1016-32 200–176 1.195 (1945) 0.701 (317) 6.378 (2886) <1 (20) 28.6 ± 3.7* 12.4 ± 0.2 (100) 1.93 gc1016-33 200–176 1.373 (2131) 0.823 (232) 6.888 (1942) <1 (20) 34.7 ± 3.6* 11.6 ± 0.2 (109) 2.15 gc1016-35 146–100 1.370 (2131) 1.660 (605) 2.734 (996) <1 (20) 183.4 ± 18.0* 11.0 ± 0.3 (59) 2.03 gc1016-36 146–100 1.368 (2131) 3.073 (1050) 3.626 (1239) 7 (20) 226.6 ± 13.7* 11.1 ± 0.1 (103) 1.49 gc1016-37 251–200 1.365 (2131) 1.975 (640) 2.734 (886) <1 (20) 189.4 ± 19.3* 11.3 ± 0.2 (104) 2.10 gc1016-38 251–200 1.362 (2131) 2.561 (1109) 3.167 (1371) <1 (20) 224.3 ± 16.2* 11.0 ± 0.2 (106) 1.84 gc1016-39 251–200 1.360 (2131) 3.613 (973) 5.101 (1374) <1 (17) 169.9 ± 18.1* 11.9 ± 0.2 (100) 1.64 gc1016-41 112–65 1.357 (2131) 0.736 (237) 1.695 (546) <1 (16) 128.7 ± 16.1* 11.7 ± 0.4 (40) 2.22 gc1016-42 488–359 1.355 (2131) 5.894 (2422) 11.576 (4757) <1 (20) 141.7 ± 15.6* 11.5 ± 0.2 (110) 1.77 gc1016-43 488–444 1.352 (2131) 4.992 (1310) 6.360 (1669) <1 (20) 202.8 ± 12.5* 11.7 ± 0.1 (123) 1.45 gc1016-44 200–146 1.349 (2131) 4.736 (1228) 5.357 (1389) <1 (19) 224.9 ± 20.1* 12.3 ± 0.2 (103) 1.55 gc1016-45 488–444 1.347 (2131) 4.937 (2122) 6.414 (2757) 9 (20) 199.7 ± 9.2* 10.4 ± 0.2 (121) 1.72 gc1016-46 >400 1.344 (2131) 7.544 (2051) 8.298 (2256) 17 (20) 235.7 ± 9.9 12.4 ± 0.1 (108) 1.50 gc1016-47 >400 1.341 (2131) 3.651 (1011) 5.580 (1545) <1 (20) 168.9 ± 10.7* 12.2 ± 0.1 (106) 1.54 gc1016-48 >400 1.339 (2131) 0.532 (272) 2.286 (1168) 10 (20) 60.8 ± 5.2* 11.2 ± 0.2 (104) 1.87 gc1016-49 359–299 1.234 (1997) 2.285 (1874) 2.278 (1868) <1 (20) 229.8 ± 21.7* 12.9 ± 0.1 (100) 1.29 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 102 of 154 www.geusbul let in.org appendix 1.3 afta data. also available online in supplementary file s4 (continued) sample number stratigraphic age a (ma) ρd b (106 tracks /cm2) ρs b (106 tracks /cm2) ρi b (106 tracks /cm2) p(χ2) c (number of grains) fission-track age d (ma) mean track length e (μm) (number of lengths) standard deviation f (μm) gc1016-50 359–299 1.240 (1997) 2.252 (1859) 1.977 (1632) <1 (20) 256.4 ± 14.8* 12.6 ± 0.1 (113) 1.33 gc1016-51 359–299 1.245 (1997) 0.927 (805) 1.214 (1054) <1 (20) 189.8 ± 18.8* 12.2 ± 0.2 (94) 1.60 gc1016-52 359–299 1.251 (1997) 2.662 (1920) 2.456 (1771) <1 (20) 236.0 ± 18.6* 12.1 ± 0.2 (101) 1.54 gc1016-53 359–299 1.256 (1997) 3.326 (3106) 3.195 (2984) <1 (20) 247.2 ± 19.0* 12.4 ± 0.1 (100) 1.39 gc1016-54 359–299 1.261 (1997) 2.395 (1842) 2.475 (1903) 16 (20) 225.5 ± 10.6* 12.2 ± 0.2 (99) 1.63 gc1016-55 359–299 1.267 (1997) 1.482 (1330) 2.094 (1879) <1 (20) 174.8 ± 13.9* 12.3 ± 0.2 (99) 1.58 gc1016-57 >400 1.272 (1997) 1.104 (708) 1.528 (980) <1 (15) 174.7 ± 14.0* 12.2 ± 0.2 (59) 1.58 gc1016-58 >400 1.277 (1997) 2.406 (2476) 2.960 (3046) <1 (20) 190.3 ± 11.1* 11.6 ± 0.2 (101) 1.53 gc1016-59 >400 1.283 (1997) 2.788 (1660) 6.351 (3781) <1 (20) 108.8 ± 8.0* 11.6 ± 0.2 (109) 2.01 gc1016-60 >400 1.288 (1997) 2.243 (2003) 5.544 (4951) <1 (20) 98.2 ± 6.2* 12.1 ± 0.2 (105) 1.86 gc1016-61 >400 1.294 (1997) 1.205 (711) 4.935 (2913) <1 (20) 60.6 ± 6.4* 12.4 ± 0.2 (102) 2.22 gc1016-62 >400 1.299 (1997) 0.558 (526) 1.679 (1582) <1 (20) 72.0 ± 8.5* 12.5 ± 0.2 (86) 2.18 gc1016-63 >400 1.348 (2106) 2.795 (1261) 4.291 (1936) 5 (20) 172.1 ± 9.5* 11.4 ± 0.2 (127) 2.02 gc1016-65 >400 1.346 (2106) 0.473 (356) 3.361 (2532) 13 (20) 37.3 ± 2.7* 12.7 ± 0.2 (104) 2.43 gc1016-66 299–251 1.345 (2106) 1.539 (1264) 1.883 (1546) <1 (20) 208.2 ± 17.0* 11.3 ± 0.2 (106) 1.68 gc1016-67 251–200 1.343 (2106) 2.903 (1257) 2.520 (1091) <1 (20) 291.2 ± 20.1* 11.8 ± 0.1 (106) 1.36 gc1016-68 251–200 1.342 (2106) 2.170 (1521) 2.033 (1425) 8 (20) 276.2 ± 15.2* 12.1 ± 0.1 (125) 1.58 gc1016-69 251–200 1.340 (2106) 1.514 (946) 1.818 (1136) 12 (20) 216.3 ± 13.3* 11.3 ± 0.2 (109) 2.41 gc1016-70 359–299 1.339 (2106) 2.344 (1326) 2.395 (1355) 8 (20) 249.8 ± 14.5* 11.9 ± 0.2 (131) 1.87 gc1016-72 >400 1.337 (2106) 0.770 (500) 1.012 (657) 20 (20) 196.9 ± 13.0 11.5 ± 0.2 (101) 2.16 gc1016-73 >400 1.336 (2106) 0.353 (308) 2.437 (2124) <1 (20) 39.2 ± 3.6* 12.6 ± 0.3 (100) 2.65 gc1077-1 >400 1.102 (1737) 1.866 (1038) 2.191 (1219) 53 (20) 176.1 ± 9.0 13.9 ± 0.1 (101) 1.28 gc1077-3 >400 1.106 (1737) 0.368 (278) 0.701 (530) 31 (20) 109.4 ± 8.7 13.3 ± 0.2 (65) 1.82 gc1077-4 >400 1.109 (1737) 5.724 (3040) 5.411 (2874) <1 (20) 214.2 ± 14.4* 13.1 ± 0.1 (103) 1.38 gc1077-5 >400 1.112 (1737) 4.511 (1740) 5.669 (2187) 1 (20) 168.0 ± 8.7* 12.3 ± 0.2 (103) 1.71 gc1077-6 >400 1.115 (1737) 1.028 (884) 1.325 (1139) <1 (20) 160.7 ± 11.4* 13.2 ± 0.1 (100) 1.19 gc1077-8 >400 1.119 (1737) 1.470 (1795) 1.764 (2154) <1 (20) 180.8 ± 12.6* 13.4 ± 0.1 (112) 1.58 gc1077-9 >400 1.122 (1737) 1.338 (1608) 1.419 (1706) <1 (20) 197.6 ± 11.2* 12.6 ± 0.2 (106) 1.66 gc1077-10 >400 1.002 (1612) 0.380 (352) 0.335 (310) 78 (20) 219.7 ± 18.4 12.5 ± 0.1 (109) 1.52 gc1077-11 >400 1.005 (1612) 1.711 (869) 1.447 (735) 7 (20) 231.3 ± 16.3* 12.8 ± 0.2 (112) 2.43 gc1077-12 130–125 1.008 (1612) 2.597 (1332) 1.965 (1008) 61 (20) 256.6 ± 13.4 11.8 ± 0.2 (105) 1.77 gc1077-13 130–125 1.012 (1612) 1.384 (499) 1.348 (486) <1 (20) 203.4 ± 20.8* 11.3 ± 0.2 (104) 2.39 gc1077-15 176–161 1.015 (1612) 3.047 (1415) 2.397 (1113) 20 (20) 248.6 ± 12.6 12.2 ± 0.2 (105) 1.66 gc1077-16 416–398 1.018 (1612) 2.230 (1301) 1.611 (940) 48 (20) 271.0 ± 14.4 12.2 ± 0.1 (107) 1.41 gc1077-17 165–161 1.021 (1612) 2.547 (723) 2.304 (654) 33 (20) 218.1 ± 13.6 11.7 ± 0.2 (102) 1.68 gc1077-18 800 1.452 (2181) 1.510 (496) 2.350 (772) 18 (20) 175.0 ± 11.1 12.5 ± 0.2 (31) 1.37 gc1077-20 800 1.443 (2181) 0.600 (153) 1.310 (334) 9 (19) 120.9 ± 15.5* 12.5 ± 0.4 (9) 1.15 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 103 of 154 www.geusbul let in.org appendix 1.3 afta data. also available online in supplementary file s4 (continued) sample number stratigraphic age a (ma) ρd b (106 tracks /cm2) ρs b (106 tracks /cm2) ρi b (106 tracks /cm2) p(χ2) c (number of grains) fission-track age d (ma) mean track length e (μm) (number of lengths) standard deviation f (μm) gc1077-21 800 1.434 (2181) 0.588 (77) 0.909 (119) 17 (6) 174.2 ± 25.9 12.3 ± 0.8 (7) 2.00 gc1077-22 416–398 1.426 (2181) 4.937 (1631) 6.792 (2244) <1 (20) 198.6 ± 15.0* 12.2 ± 0.1 (102) 1.15 gc1077-23 416–398 1.417 (2181) 1.635 (782) 2.287 (1094) <1 (20) 187.2 ± 16.5* 12.1 ± 0.2 (70) 1.48 gc1077-24 416–398 1.408 (2181) 1.910 (1106) 2.344 (1357) <1 (20) 206.5 ± 15.9* 11.8 ± 0.2 (100) 1.87 gc1077-25 416–398 1.399 (2181) 6.523 (3904) 8.966 (5366) <1 (20) 190.3 ± 8.2* 12.1 ± 0.2 (101) 1.80 gc1077-26 416–398 1.391 (2181) 3.155 (1892) 4.537 (2721) <1 (20) 177.7 ± 9.8* 11.6 ± 0.2 (51) 1.47 gc1077-27 416–398 1.382 (2181) 1.159 (808) 1.651 (1151) <1 (20) 180.0 ± 14.0* 13.2 ± 0.2 (62) 1.61 gc1077-28 416–398 1.373 (2181) 0.662 (570) 0.981 (845) <1 (20) 164.7 ± 14.4* 13.0 ± 0.2 (42) 1.48 gc1077-29 416–398 1.364 (2181) 0.764 (563) 1.411 (1040) <1 (20) 146.7 ± 12.5* 13.1 ± 0.2 (41) 1.44 gc1077-30 416–398 1.356 (2181) 1.022 (632) 1.495 (925) <1 (20) 178.3 ± 14.4* 13.4 ± 0.2 (29) 1.11 gc1077-31 416–398 1.347 (2181) 0.660 (541) 1.267 (1039) <1 (20) 125.2 ± 14.0* 12.0 ± 0.3 (27) 1.54 gc1077-32 416–398 1.338 (2181) 3.288 (2034) 5.359 (3315) <1 (20) 151.9 ± 9.0* 12.9 ± 0.2 (98) 1.61 gc1077-33 416–398 1.329 (2181) 0.305 (281) 0.435 (401) 13 (20) 171.1 ± 17.2* 12.1 ± 0.2 (53) 1.38 gc1077-34 416–398 1.321 (2181) 0.168 (168) 0.380 (381) 80 (20) 109.8 ± 10.6 11.8 ± 0.2 (45) 1.28 gc1077-35 416–398 1.444 (2283) 1.049 (494) 1.897 (893) 77 (20) 155.1 ± 9.7 11.0 ± 0.2 (106) 2.23 gc1077-36 416–398 1.445 (2283) 0.048 (31) 0.349 (226) 81 (20) 38.8 ± 7.5 11.4 ± 0.4 (45) 2.58 gc1077-37 165–156 1.446 (2283) 1.748 (99) 2.048 (116) 19 (4) 238.0 ± 33.2 12.3 ± 0.4 (22) 1.65 gc1077-38 172–165 1.447 (2283) 1.621 (514) 2.478 (786) <1 (20) 165.2 ± 20.5* 10.3 ± 0.2 (113) 2.47 gc1077-39 190–176 1.447 (2283) 0.399 (205) 2.466 (1268) <1 (20) 49.2 ± 6.3* 11.8 ± 0.2 (112) 2.19 gc1077-40 66–34 1.448 (2283) 1.883 (1344) 2.187 (1561) <1 (20) 240.4 ± 11.2 12.3 ± 0.2 (112) 2.20 gc1077-41 540 1.273 (2005) 4.517 (3633) 5.822 (4682) <1 (20) 182.2 ± 8.1* 11.9 ± 0.1 (100) 1.27 gc1077-42 444–416 1.273 (2005) 0.075 (62) 1.269 (1051) 1 (19) 13.4 ± 2.8* 13.9 ± 1.1 (3) 1.92 gc1077-43 >400 1.274 (2005) 0.073 (56) 0.734 (563) 75 (19) 24.1 ± 3.4 13.7 ± 0.4 (3) 0.66 gc1077-44 >400 1.274 (2005) 0.355 (146) 3.648 (1499) <1 (20) 23.8 ± 3.1* 12.3 ± 0.2 (101) 1.78 gc1077-45 299–251 1.274 (2005) 0.336 (37) 2.606 (287) 4 (10) 32.2 ± 7.8* 10.5 ± 1.4 (2) 1.98 gc1077-46 251–200 1.275 (2005) 0.593 (421) 3.931 (2793) <1 (20) 37.4 ± 3.9* 13.0 ± 0.1 (100) 1.47 gc1077-47 359–299 1.275 (2005) 0.968 (426) 2.055 (904) <1 (12) 107.1 ± 16.1* 12.7 ± 0.4 (20) 1.69 gc1077-49 >400 1.275 (2005) 1.117 (388) 2.058 (715) <1 (15) 113.2 ± 18.6* 12.7 ± 0.5 (8) 1.37 gc1077-50 359–299 1.275 (2005) 0.743 (416) 3.839 (2150) <1 (20) 42.9 ± 6.8* 11.1 ± 0.3 (71) 2.22 gc1077-51 416–359 1.276 (2005) 0.244 (238) 2.187 (2137) <1 (19) 32.7 ± 6.8* 11.9 ± 0.2 (65) 1.81 gc1077-52 36 1.276 (2005) 0.043 (36) 0.291 (241) 97 (20) 36.2 ± 6.5 13.5 ± 0.5 (10) 1.54 gc1077-54 36 1.276 (2005) 0.099 (62) 0.799 (500) <1 (20) 32.6 ± 7.5* 11.9 ± 0.7 (4) 1.37 gc1077-55 36 1.277 (2005) 0.021 (19) 0.543 (496) 78 (20) 9.3 ± 2.2 12.1 ± 0.4 (4) 0.76 gc1077-56 540 1.336 (2074) 4.334 (150) 3.583 (124) 70 (3) 309.8 ± 38.7 11.9 ± 0.2 (107) 1.79 gc1077-57 540 1.338 (2074) 0.321 (243) 0.524 (397) 80 (20) 158.9 ± 13.7 12.7 ± 0.2 (101) 1.54 gc1077-58 >400 1.224 (1868) 0.591 (569) 1.330 (1281) <1 (20) 103.3 ± 9.1* 12.6 ± 0.2 (76) 1.61 gc1077-59 >400 1.221 (1868) 1.226 (1356) 2.220 (2456) <1 (20) 124.1 ± 11.0* 12.8 ± 0.1 (102) 1.43 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 104 of 154 www.geusbul let in.org appendix 1.3 afta data. also available online in supplementary file s4 (continued) sample number stratigraphic age a (ma) ρd b (106 tracks /cm2) ρs b (106 tracks /cm2) ρi b (106 tracks /cm2) p(χ2) c (number of grains) fission-track age d (ma) mean track length e (μm) (number of lengths) standard deviation f (μm) gc1077-60 540 1.217 (1868) 2.258 (1607) 3.983 (2835) <1 (20) 131.3 ± 8.3* 12.3 ± 0.2 (101) 1,55 gc1077-61 >540 1.213 (1868) 3.857 (1245) 6.852 (2212) <1 (20) 131.1 ± 7.9* 11.9 ± 0.2 (68) 2,02 gc1077-62 >400 1.209 (1868) 1.713 (1674) 2.520 (2463) <1 (20) 154.6 ± 9.6* 12.2 ± 0.2 (101) 1.74 gc1077-63 >540 1.206 (1868) 3.448 (3320) 5.326 (5128) <1 (20) 145.4 ± 7.0* 12.6 ± 0.1 (103) 1.49 gc1077-64 >540 1.202 (1868) 0.632 (730) 1.146 (1324) <1 (20) 118.4 ± 11.2* 12.6 ± 0.2 (95) 2.09 gc1077-65 >540 1.198 (1868) 2.322 (2712) 3.469 (4052) <1 (20) 147.3 ± 8.6* 12.5 ± 0.2 (102) 1.59 gc1077-66 >540 1.195 (1868) 1.575 (1764) 2.465 (2761) <1 (20) 148.0 ± 10.2* 12.5 ± 0.2 (101) 1.71 gc1077-67 540 1.191 (1868) 1.798 (1418) 3.701 (2918) <1 (20) 113.1 ± 8.5* 12.3 ± 0.2 (101) 1.58 gc1077-68 >400 1.187 (1868) 1.893 (1176) 3.151 (1957) 15 (20) 133.2 ± 7.0* 12.6 ± 0.2 (106) 1.70 gc1077-69 444–416 1.184 (1868) 6.032 (4984) 7.199 (5948) <1 (20) 180.6 ± 8.9* 12.7 ± 0.1 (102) 1.45 gc1077-70 540 1.180 (1868) 4.452 (2317) 4.602 (2395) <1 (20) 227.8 ± 15.9* 13.7 ± 0.1 (103) 1.48 gc1077-71 >540 1.176 (1868) 1.948 (2131) 2.937 (3212) <1 (20) 143.6 ± 9.6* 13.4 ± 0.1 (100) 1.48 gc1077-72 66–50 1.173 (1868) 0.892 (547) 3.872 (2373) <1 (20) 50.6 ± 4.0* 13.3 ± 0.3 (32) 1.45 gc1077-73 146–100 1.169 (1868) 0.702 (471) 2.992 (2007) <1 (20) 54.9 ± 6.8* 12.7 ± 0.2 (103) 1.55 gc1077-74 416–398 1.165 (1868) 0.825 (792) 5.558 (5337) <1 (19) 34.0 ± 2.6* 11.7 ± 0.2 (101) 1.81 gc1077-75 146–66 1.162 (1868) 0.591 (515) 3.518 (3064) <1 (20) 38.4 ± 3.4* 13.3 ± 0.1 (103) 1.33 gc1077-76 228–200 1.158 (1868) 0.350 (355) 2.157 (2191) <1 (20) 35.7 ± 3.8* 11.8 ± 0.2 (100) 2.32 gc1077-77 161–146 1.154 (1868) 0.989 (739) 2.561 (1913) <1 (20) 79.2 ± 9.3* 10.0 ± 0.3 (78) 2.31 gc1077-78 204–189 1.285 (1938) 0.334 (315) 2.160 (2039) <1 (20) 41.5 ± 4.8* 12.1 ± 0.3 (47) 2.31 gc1077-79 251–200 1.280 (1938) 0.652 (312) 3.427 (1639) <1 (18) 48.5 ± 6.2* 12.5 ± 0.2 (36) 1.34 gc1077-80 145–65 1.275 (1938) 0.636 (366) 3.565 (2053) <1 (20) 45.3 ± 4.4* 13.6 ± 0.3 (34) 1.58 gc1077-81 146–66 1.270 (1938) 2.063 (1259) 2.559 (1562) 2 (20) 190.3 ± 11.3* 11.3 ± 0.1 (105) 1.41 gc1077-82 416–398 1.265 (1938) 2.616 (1707) 3.975 (2594) <1 (20) 155.7 ± 9.8* 10.4 ± 0.2 (107) 1.94 gc1077-83 416–398 1.260 (1938) 1.403 (989) 2.168 (1528) <1 (20) 152.7 ± 9.8* 12.1 ± 0.2 (100) 1.60 gc1077-84 146–100 1.255 (1938) 1.424 (1001) 2.596 (1825) <1 (20) 128.3 ± 10.6* 12.8 ± 0.2 (104) 1.64 gc1077-85 55 1.250 (1938) 1.299 (242) 2.292 (427) <1 (12) 150.0 ± 24.2* 13.6 ± 0.5 (10) 1.60 gc1077-86 >400 1.245 (1938) 1.749 (252) 1.624 (234) 22 (5) 250.0 ± 23.7 11.7 ± 0.5 (9) 1.45 gc1077-87 251–200 1.239 (1938) 1.794 (568) 2.625 (831) <1 (20) 159.6 ± 16.7* 11.1 ± 0.3 (31) 1.59 gc1077-88 >400 1.234 (1938) 0.499 (529) 3.064 (3251) <1 (20) 38.6 ± 2.7* 12.8 ± 0.2 (103) 1.98 gc1077-89 146-100 1.229 (1938) 1.759 (1348) 1.900 (1456) <1 (20) 218.6 ± 16.0* 13.0 ± 0.1 (104) 1.40 gc1077-90 >400 1.224 (1938) 0.716 (454) 2.868 (1819) <1 (18) 59.4 ± 5.1* 13.4 ± 0.2 (103) 1.78 gc1077-91 >400 1.219 (1938) 1.479 (941) 1.987 (1264) <1 (20) 178.5 ± 13.6* 13.0 ± 0.2 (107) 1.71 gc1077-92 >400 1.214 (1938) 1.512 (922) 2.852 (1739) <1 (20) 122.6 ± 9.5* 12.7 ± 0.2 (103) 1.69 gc1077-93 >400 1.209 (1938) 2.606 (1950) 2.614 (1956) <1 (20) 227.7 ± 14.7* 12.7 ± 0.1 (100) 1.21 gc1077-94 >400 1.204 (1938) 3.022 (1339) 3.282 (1454) <1 (20) 211.5 ± 13.3* 12.3 ± 0.2 (100) 1.59 gc1077-95 >400 1.199 (1938) 2.839 (1981) 2.940 (2052) 5 (20) 217.5 ± 10.4* 12.9 ± 0.1 (103) 1.07 gc1077-96 65–56 1.194 (1938) 1.327 (1322) 1.679 (1673) <1 (20) 167.3 ± 21.0* 13.2 ± 0.1 (100) 1.41 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 105 of 154 www.geusbul let in.org appendix 1.3 afta data. also available online in supplementary file s4 (continued) sample number stratigraphic age a (ma) ρd b (106 tracks /cm2) ρs b (106 tracks /cm2) ρi b (106 tracks /cm2) p(χ2) c (number of grains) fission-track age d (ma) mean track length e (μm) (number of lengths) standard deviation f (μm) gc1077-97 >400 1.189 (1938) 2.846 (1999) 3.543 (2488) <1 (20) 169.1 ± 14.1* 10.9 ± 0.1 (109) 1.39 gc1077-98 161–146 1.184 (1938) 0.788 (247) 1.532 (480) 19 (20) 114.8 ± 9.5 11.4 ± 0.5 (20) 2.39 gc1077-99 416–398 1.190 (1977) 4.326 (1293) 7.638 (2283) <1 (20) 135.7 ± 11.2* 11.5 ± 0.2 (111) 2.34 gc1077-100 23–16 1.196 (1977) 0.063 (37) 0.608 (356) 85 (20) 24.4 ± 4.3 14.0 ± 0.3 (30) 1.79 gc1077-101 140–136 1.203 (1977) 1.251 (1101) 1.101 (969) 15 (20) 263.2 ± 16.6* 11.7 ± 0.2 (106) 2.02 gc1077-102 416–398 1.209 (1977) 2.033 (495) 2.435 (593) 4 (17) 198.6 ± 16.5* 11.5 ± 0.2 (108) 1.66 gc1077-103 416–398 1.215 (1977) 2.547 (1183) 2.332 (1083) 7 (20) 253.7 ± 14.8* 13.4 ± 0.1 (104) 1.39 gc1077-104 416–398 1.222 (1977) 1.717 (604) 1.643 (578) 48 (20) 246.0 ± 16.0 12.4 ± 0.2 (114) 2.04 gc1077-105 65–56 1.228 (1977) 1.229 (362) 1.681 (495) <1 (17) 189.3 ± 24.0* 12.8 ± 0.2 (112) 1.87 gc1104-1 >400 1.203 (1913) 0.770 (493) 0.695 (445) 54 (21) 256.7 ± 18.4 12.9 ± 0.2 (107) 1.86 gc1104-2 >400 1.212 (1913) 0.418 (301) 0.427 (307) 65 (20) 229.3 ± 19.8 12.7 ± 0.1 (107) 1.35 gc1104-3 >400 1.220 (1913) 1.229 (785) 1.335 (853) 30 (20) 216.9 ± 12.5 13.0 ± 0.1 (114) 1.32 gc1104-4 >400 1.229 (1913) 1.537 (412) 1.537 (412) 27 (20) 236.9 ± 17.9 10.9 ± 0.2 (105) 2.08 gc1104-5 >400 1.237 (1913) 2.736 (1157) 3.133 (1325) 2 (20) 208.8 ± 13.0* 12.2 ± 0.1 (151) 1.53 gc1104-6 >400 1.245 (1913) 4.138 (1440) 4.121 (1434) 12 (20) 241.1 ± 11.5 12.2 ± 0.2 (106) 1.64 gc1104-7 >400 1.254 (1913) 3.160 (1197) 3.519 (1333) 11 (20) 217.5 ± 10.8 11.9 ± 0.1 (115) 1.51 gc1104-8 >400 1.262 (1913) 1.214 (964) 1.316 (1045) 93 (20) 224.8 ± 12.0 11.6 ± 0.2 (111) 1.80 gc1104-9 >400 1.271 (1913) 2.159 (1405) 1.979 (1288) 50 (20) 266.7 ± 13.0 12.8 ± 0.1 (104) 1.41 gc1104-10 >400 1.279 (1913) 0.378 (74) 2.412 (472) <1 (10) 41.0 ± 10.0* 11.9 ± 0.3 (48) 2.28 gc1104-12 >400 1.287 (1913) 2.618 (1099) 3.450 (1448) 4 (20) 189.0 ± 11.0* 11.6 ± 0.2 (103) 1.87 gc1104-13 >400 1.296 (1913) 2.589 (958) 4.262 (1577) <1 (19) 156.0 ± 12.2* 11.2 ± 0.2 (116) 1.85 gc1104-14 >400 1.145 (1820) 2.144 (1946) 3.567 (3237) <1 (20) 129.7 ± 6.2* 12.4 ± 0.2 (104) 1.59 gc1104-15 83.5–70.6 1.147 (1820) 0.343 (308) 1.149 (1032) <1 (20) 63.9 ± 9.4* 13.4 ± 0.2 (43) 1.35 gc1104-16 251–200 1.148 (1820) 0.696 (507) 3.837 (2794) <1 (20) 42.0 ± 5.6* 12.2 ± 0.2 (86) 1.78 gc1104-17 200–146 1.150 (1820) 0.310 (74) 1.866 (445) 75 (14) 36.3 ± 4.7 12.0 ± 1.6 (2) 2.31 gc1104-18 199.6–145.5 1.152 (1820) 0.590 (244) 4.820 (1993) <1 (20) 26.3 ± 2.8* 14.1 ± 0.2 (33) 1.21 gc1104-19 251–199 1.153 (1820) 1.051 (753) 5.281 (3782) <1 (20) 47.7 ± 5.3* 12.6 ± 0.2 (100) 1.59 gc1104-20 >400 1.155 (1820) 2.540 (1512) 3.564 (2122) <1 (20) 152.5 ± 10.9* 11.8 ± 0.1 (100) 1.34 gc1104-21 >400 1.156 (1820) 3.385 (2075) 4.813 (2950) <1 (20) 151.0 ± 8.6* 11.0 ± 0.1 (102) 1.48 gc1104-22 >400 1.158 (1820) 3.158 (924) 4.238 (1240) <1 (20) 170.7 ± 18.9* 12.0 ± 0.1 (89) 1.37 gc1104-23 140.2–136.4 1.159 (1820) 1.416 (484) 1.823 (623) 6 (20) 167.2 ± 13.6* 11.7 ± 0.2 (101) 1.59 gc1104-24 85.8–83.5 1.477 (2262) 1.329 (603) 2.133 (968) <1 (20) 190.0 ± 18.0* 10.9 ± 0.2 (104) 2.14 a all numerical values for stratigraphic ages assigned following gradstein et al. (2012). b ρs = spontaneous track density. ρi = induced track density. ρd = glass dosimeter track density. numbers in parentheses show the number of tracks counted in determining track densities. c probability that all single-grain ages belong to a single population (galbraith 2005). d central age (galbraith 2005), used for samples containing a significant spread in single-grain ages (p(χ2) < 5%) is denoted with *, otherwise the pooled age is quoted. all errors quoted at ± 1σ. e all analytical details are as provided in supplementary file s1, s2, s3. f standard deviation of the track-length distribution. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 106 of 154 www.geusbul let in.org a pp en di x 1. 4 th er m al h is to ry s ol ut io ns . a ls o av ai la bl e on lin e in s up pl em en ta ry f ile s 4 sa m pl e nu m be r / re gi on vt el ev at io n (m a .s .l. ) st ra tig ra ph ic su bd iv is io n st ra tigr ap hi c ag e (m a) pa la eo te m p. (° c) o ns et o f co ol in g (m a) ev en t pa la eo te m p. (° c) o ns et o f co ol in g (m a) ev en t pa la eo te m p. (° c) o ns et o f co ol in g (m a) ev en t re gi on 1 : s to re k o ld ew ey r eg io n 1a : g er m an ia l an d g c1 01 619 1 pr ec am br ./c al ed . >4 00 >1 10 26 0– 18 0 c1 75 –9 5 18 0– 45 c4 30 –6 5 40 –0 c8 g c1 10 41 28 0 pr ec am br ./c al ed . >4 00 >1 00 b >3 00 b c0 95 –1 05 29 5– 18 5 c1 50 –6 5 35 –0 c8 g c1 07 710 20 0 pr ec am br ./c al ed . >4 00 >1 05 >2 50 c0 10 0– 10 5 27 0– 15 0 c1 60 –7 0 50 –0 c8 g c1 07 79 20 pr ec am br ./c al ed . >4 00 >1 00 26 5– 10 0 c1 75 –9 5 18 5– 75 c4 40 –7 0 60 –0 c8 g c1 10 42 19 0 pr ec am br ./c al ed . >4 00 >1 00 32 0– 19 0 c1 65 –7 5 50 –0 c8 g c1 07 711 51 0 pr ec am br ./c al ed . >4 00 >1 05 >2 50 c0 10 0– 10 5 27 0– 18 0 c1 65 –7 5 12 0– 50 c4 1b : s to re k ol de w ey g c1 07 713 34 ba rr em ia n 13 0– 12 5 90 –1 00 b >5 0 b c4 60 –8 0 40 –0 c8 g c1 07 712 54 ba rr em ia n 13 0– 12 5 75 –1 00 b >2 0 b c4 40 –8 0 40 –0 c8 g c1 07 710 4 52 4 d ev on ia n, e 41 6– 39 8 >1 00 b >2 20 b c0 80 –9 0 22 0– 11 5 c3 45 –7 0 65 –1 5 a c8 g c1 07 710 1 77 va la ng in ia n? 14 0– 13 6 80 –1 00 14 0– 35 c4 35 –7 5 40 –0 c8 g c1 07 710 2 77 d ev on ia n, e 41 6– 39 8 >1 00 35 0– 22 0 c0 85 –9 5 23 5– 35 c4 45 –8 0 30 –0 c8 g c1 07 783 10 7 d ev on ia n, e 41 6– 39 8 >1 00 21 0– 15 0 c2 80 –8 5 11 5– 45 c4 40 –6 5 20 –0 c8 g c1 07 710 3 66 3 d ev on ia n, e 41 6– 39 8 >1 20 32 5– 25 0 c0 65 –8 0 11 5– 45 c4 g c1 07 715 10 8 ju ra ss ic , m 16 1– 15 6 80 –1 05 b >8 0 b c4 55 –7 0 45 –0 c8 g c1 07 716 10 5 d ev on ia n, e 41 6– 39 8 >1 05 37 0– 28 0 c0 75 –9 5 28 0– 50 c4 40 –7 0 40 –0 c8 g c1 10 422 54 pr ec am br ./c al ed . >4 00 >1 00 24 5– 19 0 c1 75 –8 5 13 5– 15 c4 30 –7 0 15 –0 c8 g c1 10 423 54 va la ng in ia n 14 0– 13 6 75 –9 0 13 0– 10 c4 20 –7 5 20 –0 c8 g c1 10 420 48 pr ec am br ./c al ed . >4 00 >1 05 20 5– 16 5 c2 75 –8 5 20 –0 c8 1c : d ov e bu gt g c1 07 73 90 0 pr ec am br ./c al ed . >4 00 >1 00 ~2 00 b c1 95 –1 05 13 5– 65 c4 45 –7 5 65 –0 c8 g c1 10 43 83 0 pr ec am br ./c al ed . >4 00 >1 05 27 5– 20 0 c1 60 –7 0 55 –5 c8 g c1 07 78 77 5 pr ec am br ./c al ed . >4 00 >1 00 21 0– 17 5 c2 80 –1 00 19 0– 12 0 c3 45 –6 5 80 –2 0 a c8 g c1 07 71 18 00 pr ec am br ./c al ed . >4 00 >1 10 20 5– 16 5 c2 40 –6 5 12 5– 0 c8 g c1 07 75 34 0 pr ec am br ./c al ed . >4 00 >1 00 22 0– 18 0 c2 80 –9 0 10 5– 45 c4 30 –6 0 35 –0 c8 g c1 07 76 0 pr ec am br ./c al ed . >4 00 >1 05 20 5– 16 0 c2 55 –6 5 40 –0 c8 1d : h oc hs te tt er f or la nd e tc . g c1 07 74 10 0 pr ec am br ./c al ed . >4 00 >1 00 >2 50 c0 80 –1 00 25 0– 16 0 c2 45 –6 0 40 –0 c8 g c1 07 717 2 ca llo vi an 16 1– 15 6 80 –1 05 b >4 0 b c4 55 –8 0 35 –0 c8 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 107 of 154 www.geusbul let in.org a pp en di x 1. 4 th er m al h is to ry s ol ut io ns . a ls o av ai la bl e on lin e in s up pl em en ta ry f ile s 4 (c on tin ue d) sa m pl e nu m be r / re gi on vt el ev at io n (m a .s .l. ) st ra tig ra ph ic su bd iv is io n st ra tigr ap hi c ag e (m a) pa la eo te m p. (° c) o ns et o f co ol in g (m a) ev en t pa la eo te m p. (° c) o ns et o f co ol in g (m a) ev en t pa la eo te m p. (° c) o ns et o f co ol in g (m a) ev en t g c1 07 760 0 pr ec am br ./c al ed . >4 00 >1 15 16 5– 13 5 c3 85 –1 05 13 0– 55 c4 60 –8 0 10 –0 c8 g c1 07 761 36 0 pr ec am br ./c al ed . >4 00 >1 00 b ? c2 /c 3 90 –1 00 13 0– 75 c4 60 –7 5 20 –0 c8 re gi on 2 : ku h n ø , w o ll a st o n f o re la n d , p a ye r la n d , c la ve ri n g ø 2a : k uh n ø g c1 07 722 vt 9 76 5 d ev on ia n, e 41 6– 39 8 >1 05 26 0– 21 5 c1 70 –8 0 20 –0 a c4 g c1 07 723 vt 9 57 2 d ev on ia n, e 41 6– 39 8 >1 05 26 5– 20 5 c1 80 –8 5 12 0– 35 c4 35 –6 5 25 –0 c8 g c1 07 724 vt 9 40 0 d ev on ia n, e 41 6– 39 8 >1 00 >2 00 b c1 85 –9 5 17 5– 60 c4 50 –8 0 50 –0 c8 g c1 07 725 vt 9 22 5 d ev on ia n, e 41 6– 39 8 >1 00 >2 00 b c1 85 –9 5 17 0– 90 c4 50 –7 0 35 –0 c8 g c1 07 726 vt 9 10 0 d ev on ia n, e 41 6– 39 8 >1 00 26 0– 21 5 c1 80 –9 0 10 0– 20 c4 35 –7 0 20 –0 c8 2b : w ol la st on f or el an d g c1 07 798 50 ju ra ss ic , l 16 1– 14 6 85 –1 00 11 0– 20 c4 30 –8 0 45 –0 c8 g c1 07 799 10 0 d ev on ia n, e 41 6– 39 8 >1 05 b >1 50 b c1 10 0– 10 5 12 0– 60 c4 65 –8 0 35 –5 c8 g c1 07 772 40 2 pa le oc en e– e. e oc en e 65 –5 0 >8 5 b >3 0 b c6 60 –8 5 40 –5 c8 g c1 07 773 15 0 cr et ac eo us , e 14 6– 10 0 11 0– 12 0 65 –4 5 c5 70 –8 0 20 –1 0 c8 g c1 07 774 15 4 d ev on ia n, e 41 6– 39 8 >1 00 b ~5 5 b c5 10 0– 10 5 50 –2 5 c6 80 –9 0 10 –3 c8 g c1 07 775 34 0 cr et ac eo us 14 6– 66 >1 20 50 –3 5 c6 55 –6 5 12 –0 c8 g c1 07 710 0 81 m io ce ne . e 23 –1 6 >1 20 42 –2 0 c7 40 –9 0 23 –4 c8 g c5 22 -8 30 0 cr et ac eo us , e 14 6– 10 0 85 –1 05 b >3 5 b c4 45 –7 5 25 –0 c8 g c1 07 791 vt 16 14 46 pr ec am br ./c al ed . >4 00 >1 00 b ? c2 75 –9 0 15 5– 80 c4 35 –6 0 40 –0 c8 g c1 07 789 vt 16 50 0 cr et ac eo us , e 14 6– 10 0 60 –7 5 80 –1 7 c5 <7 5 c8 g c1 07 790 vt 16 50 0 pr ec am br ./c al ed . >4 00 >1 15 b ? c3 11 0– 11 5 70 –5 5 c5 45 –6 5 30 –5 c8 g c1 07 792 vt 16 23 5 pr ec am br ./c al ed . >4 00 >1 10 15 5– 12 0 c3 95 –1 10 13 5– 90 c4 60 –7 5 40 –5 c8 g c1 01 63 65 4 cr et ac eo us , e 14 6– 10 0 60 –9 5 b >6 0 b c4 60 –7 0 37 –0 c8 g c5 22 -7 15 0 ju ra ss ic , l 16 1– 14 6 85 –9 5 70 –2 0 c6 20 –7 5 26 –0 c8 g c1 01 623 64 pa le oc en e 66 –5 6 60 –8 5 40 –5 c6 g c1 01 648 28 pr ec am br ./c al ed . >4 00 >1 10 ~1 00 c4 95 –1 05 67 –3 2 c6 75 –8 5 12 –0 c8 g c1 07 788 32 7 pr ec am br ./c al ed . >4 00 >1 00 b ? c4 10 0– 10 5 38 –2 4 c6 45 –6 5 12 –0 c8 g c1 07 784 58 cr et ac eo us , e 14 6– 10 0 60 –1 00 b >1 5 b c6 55 –6 5 40 –0 c8 g c1 07 710 5 42 0 pa le oc en e 66 –5 6 45 –8 5 b >5 b c8 2c : p ay er l an d g c1 07 770 11 50 pr ec am br ./c al ed . >4 00 >1 30 25 5– 20 0 c1 70 –9 0 13 5– 50 c4 <8 0 80 –0 c8 g c1 07 749 41 3 pr ec am br ./c al ed . >4 00 >1 15 b ? c1 /c 2 10 5– 11 5 15 0– 10 0 c3 40 –9 0 90 –0 c8 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 108 of 154 www.geusbul let in.org a pp en di x 1. 4 th er m al h is to ry s ol ut io ns . a ls o av ai la bl e on lin e in s up pl em en ta ry f ile s 4 (c on tin ue d) sa m pl e nu m be r / re gi on vt el ev at io n (m a .s .l. ) st ra tig ra ph ic su bd iv is io n st ra tigr ap hi c ag e (m a) pa la eo te m p. (° c) o ns et o f co ol in g (m a) ev en t pa la eo te m p. (° c) o ns et o f co ol in g (m a) ev en t pa la eo te m p. (° c) o ns et o f co ol in g (m a) ev en t g c1 07 747 12 0 ca rb on ife ro us 35 9– 29 9 >1 05 15 0– 11 0 c3 55 –8 0 55 –0 c8 g c1 07 727 vt 10 12 61 d ev on ia n, e 41 6– 39 8 >1 00 23 0– 17 0 c2 70 –9 5 18 0– 80 c3 25 –6 0 80 –0 c8 g c1 07 728 vt 10 93 9 d ev on ia n, e 41 6– 39 8 >1 00 22 0– 16 0 c2 55 –7 0 60 –0 c8 g c1 07 729 vt 10 64 0 d ev on ia n, e 41 6– 39 8 >1 00 18 5– 13 0 c2 55 –7 0 50 –5 c8 g c1 07 730 vt 10 47 9 d ev on ia n, e 41 6– 39 8 >1 15 22 5– 17 0 c2 50 –7 5 90 –0 c8 g c1 07 731 vt 10 17 1 d ev on ia n, e 41 6– 39 8 >1 00 19 5– 13 5 c2 75 –8 5 50 –5 c8 g c1 07 732 vt 10 32 d ev on ia n, e 41 6– 39 8 >1 00 19 5– 14 5 c2 80 –9 0 14 0– 75 c4 40 –6 0 25 –0 c9 2d : c la ve ri ng ø g c1 01 620 1 pr ec am br ./c al ed . >4 00 >1 10 23 0– 17 5 c2 75 –9 5 17 0– 60 c3 40 –6 5 25 –0 c8 g c1 07 780 45 7 cr et ac eo us ? 14 6– 66 >1 05 60 –4 0 c5 40 –8 0 32 –2 c8 g c1 07 781 30 4 cr et ac eo us 14 6– 66 85 –1 05 b >1 5 b c6 40 –8 0 20 –0 c8 g c1 07 782 30 3 d ev on ia n, e 41 6– 39 8 >1 00 30 0– 20 0 c1 90 –9 5 70 –1 0 c6 35 –8 5 15 –0 c8 g c1 07 796 vt 17 12 33 pa le oc en e, s ub -b as al t 65 –5 5 50 –8 0 60 –8 c8 g c1 07 795 vt 17 12 33 pr ec am br ./c al ed . >4 00 >1 10 27 5– 23 0 c1 60 –7 0 40 –0 c8 g c1 07 794 vt 17 89 0 pr ec am br ./c al ed . >4 00 >1 20 27 5– 22 5 c1 85 –1 05 22 5– 11 5 c2 60 –8 0 35 –0 c8 g c1 07 793 vt 17 82 0 pr ec am br ./c al ed . >4 00 >1 25 29 5– 24 0 c1 85 –1 15 26 0– 15 5 c2 60 –7 5 25 –0 c8 g c1 07 797 vt 17 42 7 pr ec am br ./c al ed . >4 00 >1 00 27 0– 22 0 c1 85 –9 0 16 –0 c8 g c1 01 649 vt 1 84 6 ca rb on ife ro us 35 9– 29 9 10 5– 11 5 28 5– 23 5 c1 70 –9 5 23 5– 65 c3 40 –6 5 25 –0 c8 g c1 01 650 vt 1 75 5 ca rb on ife ro us 35 9– 29 9 >1 00 b >2 00 b c1 75 –1 00 32 0– 60 c3 40 –7 0 45 –0 c8 g c1 01 651 vt 1 57 4 ca rb on ife ro us 35 9– 29 9 >1 05 25 0– 20 0 c1 80 –9 0 15 5– 65 c3 45 –7 0 25 –0 c8 g c1 01 652 vt 1 49 9 ca rb on ife ro us 35 9– 29 9 >1 00 b >2 00 b c1 85 –1 00 29 0– 13 5 c3 60 –7 5 45 –0 c8 g c1 01 653 vt 1 50 1 ca rb on ife ro us 35 9– 29 9 >1 00 b >2 00 b c1 85 –1 00 23 5– 13 5 c3 55 –7 0 15 –0 c8 g c1 01 654 vt 1 24 8 ca rb on ife ro us 35 9– 29 9 >1 00 b >2 00 b c1 80 –9 5 22 0– 10 5 c3 55 –7 0 35 –0 c8 g c1 01 655 vt 1 52 ca rb on ife ro us 35 9– 29 9 >1 00 b >2 00 b c1 95 –1 05 20 0– 14 0 c3 65 –7 5 30 –0 c8 g c1 07 786 77 8 pr ec am br ./c al ed . >4 00 >1 05 40 0– 25 0 c0 70 –9 0 13 0– 0 c6 g c1 07 787 re gi on 3 : h o ld w it h h o pe , h u d so n l a n d , y m er ø , g eo g ra ph ic a l so ci et y ø 3a : h w h 77 8 tr ia ss ic 25 1– 20 0 85 –9 5 40 –0 c6 g c1 07 740 73 9 eo ce ne , e 55 65 –8 5 b >2 0 b c6 g c1 01 668 vt 2 32 3 tr ia ss ic 25 1– 20 0 80 –1 00 b >5 5 b c3 45 –8 0 50 –0 c8 g c1 01 667 vt 2 29 2 tr ia ss ic 25 1– 20 0 80 –1 00 b >2 5 b c3 40 –8 0 20 –0 c8 g c1 01 669 vt 2 94 tr ia ss ic 25 1– 20 0 95 –1 00 17 5– 80 c3 60 –8 0 45 –5 c8 g c1 01 666 vt 2 0 pe rm ia n 29 9– 25 1 85 –1 00 22 0– 20 c3 50 –8 5 20 –0 c8 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 109 of 154 www.geusbul let in.org a pp en di x 1. 4 th er m al h is to ry s ol ut io ns . a ls o av ai la bl e on lin e in s up pl em en ta ry f ile s 4 (c on tin ue d) sa m pl e nu m be r / re gi on vt el ev at io n (m a .s .l. ) st ra tig ra ph ic su bd iv is io n st ra tigr ap hi c ag e (m a) pa la eo te m p. (° c) o ns et o f co ol in g (m a) ev en t pa la eo te m p. (° c) o ns et o f co ol in g (m a) ev en t pa la eo te m p. (° c) o ns et o f co ol in g (m a) ev en t g c1 07 779 15 0 tr ia ss ic 25 1– 20 0 11 5– 12 5 62 –4 2 c5 60 –8 0 9– 0 c8 g c1 01 639 75 1 tr ia ss ic 25 1– 20 0 85 –1 05 22 0– 75 c4 60 –8 0 40 –1 0 c6 g c1 01 638 38 1 tr ia ss ic 25 1– 20 0 80 –9 0 12 0– 20 c4 40 –8 0 20 –0 c8 g c1 10 424 -5 sa nt on ia n 86 –8 3 90 –1 00 >4 0 c5 60 –8 0 30 –0 c8 g c1 01 641 21 cr et ac eo us , l 11 2– 65 60 –1 00 b ~3 6 b c6 60 –8 0 37 –2 c8 3b : h ud so n la nd g c1 01 657 vt 4 12 08 pr ec am br ./c al ed . >4 00 >1 00 24 0– 17 5 c1 75 –8 0 60 –1 0 c6 g c1 01 658 vt 4 95 8 pr ec am br ./c al ed . >4 00 >1 05 27 0– 22 0 c1 85 –9 0 17 0– 35 c6 50 –8 0 25 –0 c8 g c1 01 659 vt 4 72 0 pr ec am br ./c al ed . >4 00 >1 05 16 5– 12 5 c3 85 –9 5 95 –2 5 c6 40 –8 0 30 –0 c8 g c1 01 660 vt 4 51 0 pr ec am br ./c al ed . >4 00 >1 00 b >1 00 b c3 95 –1 05 10 0– 50 a c6 60 –7 5 25 –5 c8 g c1 01 661 vt 4 30 5 pr ec am br ./c al ed . >4 00 >1 00 b >1 00 b c3 95 –1 05 58 –2 c6 55 –7 5 22 –5 c8 g c1 01 662 vt 4 12 3 pr ec am br ./c al ed . >4 00 >1 20 b >6 0 b c3 11 0– 12 0 55 –3 0 c6 40 –8 0 22 –0 c8 g c1 01 663 vt 4a 56 9 pr ec am br ./c al ed . >4 00 >1 05 26 0– 20 0 c1 85 –9 5 14 0– 20 c6 10 –8 5 40 –0 c8 g c1 01 665 vt 4a 21 0 pr ec am br ./c al ed . >4 00 >1 05 >6 0 c1 10 0– 10 5 40 –1 8 c6 55 –8 0 19 –7 c8 g c1 01 670 vt 3 95 9 ca rb on ife ro us 35 9– 29 9 95 –1 00 26 0– 16 5 c2 65 –8 0 40 –5 c6 g c1 01 624 vt 3 75 6 ca rb on ife ro us 35 9– 29 9 >1 10 b ? c2 85 –1 00 17 0– 10 0 c3 60 –7 0 40 –5 c8 g c1 01 626 vt 3 47 9 ca rb on ife ro us 35 9– 29 9 >1 10 b ? c2 85 –1 00 20 0– 10 0 c3 65 –8 0 35 –5 c8 g c1 01 625 vt 3 18 9 ca rb on ife ro us 35 9– 29 9 >1 25 20 0– 15 5 c2 85 –9 0 10 0– 35 c6 35 –7 5 20 –0 c8 g c1 07 785 32 0 eo ce ne , e 55 <1 00 po st –d ep c6 <8 0 20 –0 c8 3c : k ej se r fr an z jo se ph f jo rd g c1 07 746 vt 14 62 1 tr ia ss ic 25 1– 20 0 >1 20 50 –3 5 c6 55 –6 5 10 –0 c8 g c1 07 745 vt 14 45 5 pe rm ia n 29 9– 25 1 >1 05 85 –2 5 c6 70 –1 00 25 –0 c8 g c1 07 744 vt 14 40 9 pr ec am br ./c al ed . >4 00 >1 05 35 –1 7 c6 70 –8 5 10 –3 c8 g c1 07 743 vt 14 87 pr ec am br ./c al ed . >4 00 >1 00 45 –5 c6 g c1 01 616 1 d ev on ia n 41 6– 35 9 10 0– 10 5 17 0– 90 c3 70 –9 0 b 10 0– 20 b c6 70 –8 0 15 –0 c8 g c1 01 617 1 d ev on ia n 41 6– 35 9 >1 05 15 5– 13 0 c3 85 –9 5 70 –1 5 c6 50 –7 5 12 –0 c8 g c1 01 618 1 d ev on ia n 41 6– 35 9 95 –1 05 55 –1 5 c6 70 –9 0 15 –5 c8 g c1 07 742 16 0 si lu ri an 44 4– 41 6 >1 05 30 –5 a c6 g c1 07 751 10 3 d ev on ia n 41 6– 35 9 >1 10 b >4 0 b c5 10 0– 11 0 40 –1 5 c6 80 –9 5 10 –3 c8 g c1 07 750 25 5 ca rb on ife ro us 35 9– 29 9 >1 15 b >7 0 a, b c5 10 0– 10 5 70 –4 5 a c6 90 –9 5 15 –5 c8 g c1 01 637 vt 5 85 0 tr ia ss ic 25 1– 20 0 90 –1 00 17 5– 65 c4 65 –8 5 35 –0 c8 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 110 of 154 www.geusbul let in.org a pp en di x 1. 4 th er m al h is to ry s ol ut io ns . a ls o av ai la bl e on lin e in s up pl em en ta ry f ile s 4 (c on tin ue d) sa m pl e nu m be r / re gi on vt el ev at io n (m a .s .l. ) st ra tig ra ph ic su bd iv is io n st ra tigr ap hi c ag e (m a) pa la eo te m p. (° c) o ns et o f co ol in g (m a) ev en t pa la eo te m p. (° c) o ns et o f co ol in g (m a) ev en t pa la eo te m p. (° c) o ns et o f co ol in g (m a) ev en t g c1 01 636 vt 5 37 0 cr et ac eo us , e 14 6– 10 0 40 –8 5 20 –0 c8 g c1 01 635 vt 5 30 0 cr et ac eo us , e 14 6– 10 0 80 –8 5 32 –2 c8 g c1 10 415 65 6 ca m pa ni an ? 84 –7 1 >1 05 70 –4 0 c5 45 –7 5 30 –0 c8 re gi on 4 : ce n tr a l in la n d , i n la n d s o u th 4a : c en tr al in la nd g c1 07 771 95 0 pr ec am br ./c al ed . >4 00 >1 00 18 0– 12 5 c3 ~1 00 b ~1 00 b c4 60 –7 5 80 –2 5 a c8 g c1 07 766 65 0 pr ec am br ./c al ed . >4 00 >1 00 b ? c2 85 –9 5 13 0– 75 c4 45 –6 0 15 –0 c8 g c1 10 414 11 00 pr ec am br ./c al ed . >4 00 >1 05 b >1 60 b c2 10 0– 10 5 16 0– 10 5 c3 65 –7 5 25 –0 c8 g c1 07 762 21 65 pr ec am br ./c al ed . >4 00 >1 00 b ? c2 /c 3 85 –9 5 12 5– 55 c4 40 –6 5 20 –0 c8 g c1 07 765 87 0 pr ec am br ./c al ed . >4 00 >1 00 19 5– 15 0 c2 80 –9 0 13 0– 70 c4 45 –6 0 20 –0 c8 g c1 07 767 28 0 pr ec am br ./c al ed . >4 00 >1 00 14 5– 10 0 c3 85 –9 5 11 0– 55 c4 55 –7 0 20 –0 c8 g c1 01 615 1 pr ec am br ./c al ed . >4 00 >1 05 21 5– 14 5 c2 80 –1 00 16 5– 45 c4 35 –7 0 55 –0 c8 g c1 07 768 50 0 pr ec am br ./c al ed . >4 00 >1 10 17 5– 14 0 c3 80 –9 5 12 0– 65 c4 45 –6 5 20 –0 c8 g c1 07 769 21 10 si lu ri an 44 4– 41 6 90 –1 05 20 0– 13 0 c2 55 –6 5 25 –0 c8 g c1 01 68 12 0 pr ec am br ./c al ed . >4 00 >1 05 20 0– 10 0 c2 10 –1 00 14 0– 0 c8 g c1 01 69 1 d ev on ia n 41 6– 35 9 >1 20 22 0– 17 0 c2 60 –7 5 65 –0 c8 g c1 01 614 1 pr ec am br ./c al ed . >4 00 10 0– 10 5 14 5– 20 c4 30 –8 0 40 –0 c8 g c1 01 613 1 pr ec am br ./c al ed . >4 00 10 0– 10 5 28 0– 60 c4 80 –9 5 20 –0 c8 g c5 22 -3 0 d ev on ia n 41 635 9 >1 00 b >2 50 b c0 75 –9 0 23 0– 70 c4 40 –7 0 50 –0 c8 g c1 07 763 80 0 pr ec am br ./c al ed . >4 00 >1 05 18 5– 15 5 c2 10 0– 10 5 16 5– 11 0 c3 60 –8 0 40 –0 c8 g c1 07 764 93 0 pr ec am br ./c al ed . >4 00 >1 00 b ? c2 /c 3 80 –1 00 12 5– 80 c4 45 –6 5 30 –0 c8 4b : i nl an d so ut h g c1 01 65 10 0 pe rm ia n 29 9– 25 1 >1 10 28 5– 24 0 c1 75 –1 00 23 5– 25 c4 25 –7 5 40 –0 c8 g c1 07 756 0 pr ec am br ./c al ed . >4 00 95 –1 10 24 5– 80 c2 60 –8 0 25 –0 c8 g c1 07 757 0 pr ec am br ./c al ed . >4 00 >1 05 22 0– 14 0 c2 65 –8 0 50 –1 0 c8 g c1 07 758 80 0 pr ec am br ./c al ed . >4 00 >1 10 b ~1 40 b c3 90 –1 05 12 5– 75 c4 60 –7 0 25 –0 c8 g c1 07 759 50 0 pr ec am br ./c al ed . >4 00 >1 00 16 0– 12 0 c3 80 –1 00 12 5– 75 c4 50 –6 0 20 –0 c8 4c : m iln e la nd g c1 01 64 5 si lu ri an 44 0– 41 9 >1 30 20 0– 18 0 c2 80 –1 20 19 0– 40 c4 20 –8 0 60 –0 c8 g c1 01 646 18 58 pr ec am br ./c al ed . >4 00 >1 00 30 0– 23 0 c1 80 –9 0 21 5– 60 c2 30 –6 5 40 –0 c8 g c1 01 647 2 pr ec am br ./c al ed . >4 00 >1 15 23 0– 19 0 a c1 80 –9 5 13 5– 20 c4 20 –8 0 30 –0 c8 g c1 10 44 60 0 pr ec am br ./c al ed . >4 00 >1 00 b >2 70 b c0 95 –1 05 27 0– 13 0 c2 75 –8 5 20 –0 c8 g c1 10 45 12 70 pr ec am br ./c al ed . >4 00 >1 00 b >2 40 b c0 10 0– 10 5 24 0– 16 5 c2 65 –7 5 20 –0 c8 g c1 10 46 12 80 pr ec am br ./c al ed . >4 00 >1 05 32 0– 26 0 c0 80 –1 00 27 0– 50 c2 45 –7 5 35 –0 c8 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 111 of 154 www.geusbul let in.org a pp en di x 1. 4 th er m al h is to ry s ol ut io ns . a ls o av ai la bl e on lin e in s up pl em en ta ry f ile s 4 (c on tin ue d) sa m pl e nu m be r / re gi on vt el ev at io n (m a .s .l. ) st ra tig ra ph ic su bd iv is io n st ra tigr ap hi c ag e (m a) pa la eo te m p. (° c) o ns et o f co ol in g (m a) ev en t pa la eo te m p. (° c) o ns et o f co ol in g (m a) ev en t pa la eo te m p. (° c) o ns et o f co ol in g (m a) ev en t g c1 10 421 2 pr ec am br ./c al ed . >4 00 >1 15 24 5– 20 0 c1 10 0– 11 5 b 22 5– 15 0 b c2 90 –9 5 20 –0 c8 g c1 01 644 vt 8 75 9 ju ra ss ic 20 0– 14 6 60 –7 0 30 –0 c8 g c1 01 645 vt 8 45 8 o rd ov ic ia n 48 8– 44 4 >1 00 34 0– 23 0 c1 95 –1 10 28 5– 80 c2 20 –9 0 12 –0 c8 g c1 01 643 vt 8 36 2 o rd ov ic ia n 48 8– 44 4 >1 00 b ? c1 /c 2 80 –9 0 15 0– 40 c4 50 –7 5 20 –0 c8 g c1 01 642 vt 8 41 o rd ov ic ia nd ev on ia n 48 8– 35 9 >1 00 b ? c1 /c 2 80 –9 0 50 –0 c8 g c1 01 66 1 pr ec am br ./c al ed . >4 00 >1 15 27 0– 20 0 c1 80 –1 15 24 0– 90 c4 40 –7 5 80 –0 c8 g c1 01 67 2 pr ec am br ./c al ed . >4 00 >1 00 >2 00 c1 75 –1 05 21 0– 65 c4 50 –7 0 25 –0 c8 re gi on 5 : tr a il l ø , s ta u n in g a lp er , m es te rs vi g , s ch u ch er t d a l, w eg en er h a lv ø 5a : t ra ill ø g c1 10 419 10 69 tr ia ss ic 25 1– 20 0 >1 25 60 –4 5 c5 90 –1 05 43 –1 9 c6 45 –7 5 10 –0 c9 g c1 10 416 57 0 tr ia ss ic 25 1– 20 0 >1 10 60 –4 0 c5 90 –1 15 43 –1 0 c8 30 –8 0 12 –0 c9 g c1 10 417 57 0 ju ra ss ic ? 20 0– 14 6 >1 05 55 –2 5 c5 <1 05 30 –0 c9 g c1 10 418 47 1 ju ra ss ic 20 0– 14 6 >1 20 40 –2 5 c6 30 –7 0 17 –0 c9 g c5 22 -2 30 d ev on ia n 41 6– 35 9 >1 05 22 0– 12 5 c2 90 –1 00 12 5– 30 c6 60 –8 5 30 –0 c8 g c5 22 -6 15 0 ca rb on ife ro us 35 9– 29 9 >1 00 b >1 00 b c2 90 –1 00 45 –1 0 c6 20 –8 0 16 –0 c8 g c5 22 -1 50 0 tr ia ss ic 25 1– 20 0 >1 25 45 –2 5 c6 55 –8 0 22 –7 a c9 g c5 22 -1 1 30 0 ju ra ss ic , m 16 6– 16 1 >1 10 50 –1 5 c6 10 5– 11 0 14 –4 c8 70 –9 0 5– 0 c9 g c1 01 630 vt 6 10 75 ju r. –c re t. 20 0– 65 10 5– 11 0 45 –3 0 c6 80 –9 0 20 –1 0 c8 g c1 01 629 vt 6 73 0 ju r. –c re t. 20 0– 65 >1 00 55 –2 0 c6 80 –1 00 12 –3 c8 g c1 01 633 vt 6 46 0 ju ra ss ic , e 20 0– 17 6 >1 30 42 –2 2 c6 90 –1 05 13 –2 c8 g c1 01 632 vt 6 32 3 ju ra ss ic , e 20 0– 17 6 >1 25 38 –2 5 c6 80 –9 5 12 –4 c8 g c5 22 -9 14 0 ca rb on ife ro us , l 31 3– 30 3 >1 10 55 –2 0 c6 90 –1 05 15 –5 c8 g c5 22 -1 0 23 0 ju ra ss ic , m 16 6– 16 1 >1 20 55 –2 0 c6 10 0– 11 0 27 –5 c8 40 –9 0 7– 0 c9 g c5 22 -4 14 0 tr ia ss ic 25 1– 20 0 >1 25 ~3 5 c6 10 5– 11 5 28 –1 0 c8 35 –9 0 13 –0 c9 g c1 07 752 vt 15 10 18 o lig oc en e, l 36 >1 05 65 –2 0 c6 <1 05 20 –0 c8 g c1 07 754 vt 15 68 9 o lig oc en e, l 36 65 –1 00 25 –0 c8 g c1 07 755 vt 15 12 8 o lig oc en e, l 36 >1 10 22 –5 c8 55 –9 5 8– 0 c9 5b : j am es on l an d n (m es te rs vi g, w eg ne r h al vø ) g c5 22 -5 70 ca rb on ife ro us 35 9– 29 9 >1 20 24 –1 3 c7 90 –1 05 11 –5 c8 g c1 01 611 50 pe rm ia n 29 9– 25 1 11 5– 12 0 40 –9 c7 60 –1 05 9– 0 c8 g c1 01 610 1 ca rb on ife ro us 35 9– 29 9 >1 20 24 –1 7 c7 20 –9 5 15 –0 c8 g c1 03 -1 20 0 pe rm ia n, l 25 5 >1 20 30 –1 7 c7 10 0– 11 0 21 –1 2 a c8 60 –8 0 7– 0 c9 g c1 03 -2 20 0 pe rm ia n, l 25 5 >1 20 28 –1 4 c7 10 0– 11 0 20 –1 1 a c8 80 –1 00 90 –4 c9 g c1 59 -4 40 0 pe rm ia n 29 9– 25 1 >1 10 25 –1 7 c7 60 –8 0 10 –2 c8 g c1 59 -3 25 0 tr ia ss ic , e 25 1– 24 5 >1 15 27 –1 6 c7 20 –8 0 15 –0 c8 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 112 of 154 www.geusbul let in.org a pp en di x 1. 4 th er m al h is to ry s ol ut io ns . a ls o av ai la bl e on lin e in s up pl em en ta ry f ile s 4 (c on tin ue d) sa m pl e nu m be r / re gi on vt el ev at io n (m a .s .l. ) st ra tig ra ph ic su bd iv is io n st ra tigr ap hi c ag e (m a) pa la eo te m p. (° c) o ns et o f co ol in g (m a) ev en t pa la eo te m p. (° c) o ns et o f co ol in g (m a) ev en t pa la eo te m p. (° c) o ns et o f co ol in g (m a) ev en t 5c : s ch uc he rt d al g c1 03 -6 10 50 pe rm ia n, l 25 5 10 0– 10 5 20 0– 37 c5 85 –9 5 20 –4 c7 g c2 02 -2 vt 7 70 0 pe rm oca rb on ife ro us 35 9– 24 5 >1 10 65 –4 5 c5 95 –1 00 30 –1 4 c7 40 –6 5 9– 0 c9 g c2 02 -3 vt 7 82 0 pe rm ia n, l 25 5 10 5– 11 0 70 –4 5 c5 90 –9 5 14 –3 c9 g c2 02 -4 vt 7 10 00 pe rm ia n, l 26 0– 25 1 10 0– 10 5 75 –4 5 c5 85 –9 5 20 –8 c7 g c2 02 -5 vt 7 10 00 tr ia ss ic , m 24 5– 24 0 95 –1 05 10 5– 20 c5 55 –9 0 20 –0 c7 g c2 02 -6 vt 7 13 10 tr ia ss ic , m 24 5– 24 0 95 –1 05 18 5– 35 c5 80 –9 5 30 –5 c7 g c2 02 -7 vt 7 10 00 tr ia ss ic , l 23 5– 20 8 >1 20 60 –4 0 c5 75 –8 5 20 –5 c7 g c2 02 -1 6 vt 7 82 0 pe rm ia n, l 25 5 >1 15 60 –3 0 c5 10 0– 11 0 30 –1 7 c7 65 –8 0 9– 5 c9 g c1 59 -1 vt 7 15 0 pe rm ia n, l 25 5 >1 10 56 –4 0 c5 85 –1 00 40 –1 8 c7 40 –6 5 10 –0 c9 g c1 59 -2 vt 7 26 0 pe rm ia n, l 25 5 >1 15 60 –3 0 c5 10 0– 11 0 25 –1 5 c7 50 –7 0 7– 0 c9 g c2 02 -1 vt 7 70 0 pe rm ia n, l 26 0– 25 1 95 –1 00 80 –4 0 c5 55 –7 5 25 –0 c8 5d : s ta un in g a lp er n g c1 10 413 50 pr ec am br ./c al ed . >4 00 >1 00 b >1 50 b c2 80 –9 0 95 –2 0 c4 40 –7 5 15 –0 c8 g c1 10 412 20 pr ec am br ./c al ed . >4 00 >1 00 b >1 50 b c2 80 –9 0 14 0– 35 c4 40 –7 5 20 –0 c8 g c1 10 410 12 80 pr ec am br ./c al ed . >4 00 >1 20 75 –4 0 c5 95 –1 05 35 –1 1 c7 35 –8 0 14 –0 c8 g c1 07 718 vt 11 17 85 pr ec am br ./c al ed . >4 00 >1 05 24 0– 16 0 c2 60 –7 5 40 –0 c8 g c1 07 720 vt 11 37 3 pr ec am br ./c al ed . >4 00 >1 05 19 5– 10 0 c2 60 –8 0 50 –0 c8 g c1 07 721 vt 11 12 pr ec am br ./c al ed . >4 00 >1 05 32 0– 14 5 c2 60 –9 5 14 5– 0 c4 5e : s ta un in g a lp er s g c1 10 47 10 90 pr ec am br ./c al ed . >4 00 >1 00 b >2 65 b c0 90 –1 05 26 5– 16 5 c2 70 –8 0 20 –0 c8 g c1 10 48 30 0 pr ec am br ./c al ed . >4 00 >1 00 b >2 50 b c0 85 –9 5 19 5– 65 c4 60 –8 0 40 –0 c8 g c1 10 49 14 50 pr ec am br ./c al ed . >4 00 >1 10 35 0– 27 0 c0 70 –8 0 90 –2 5 c5 g c1 07 741 0 pr ec am br ./c al ed . >4 00 >1 25 25 0– 21 5 c1 80 –9 0 15 –0 c8 re gi on 6 : ja m es o n l a n d s , l iv er po o l la n d , b lo ss ev il le k ys t 6a : j am es on l an d s g c2 02 -9 90 0 ju ra ss ic , e 20 0– 17 6 10 0– 10 5 80 –4 0 c5 65 –8 0 22 –8 c8 g c1 07 737 vt 13 12 17 ju ra ss ic 20 0– 14 6 65 –8 5 12 5– 5 c6 g c1 07 738 vt 13 85 7 ju ra ss ic 20 0– 14 6 95 –1 10 16 0– 15 c6 35 –9 0 25 –0 c8 g c1 07 739 vt 13 46 9 ju ra ss ic 20 0– 14 6 >1 10 65 –4 0 c5 95 –1 05 40 –1 8 c6 55 –8 0 12 –0 c8 g c2 02 -1 2 13 0 tr ia ss ic , l 23 5– 20 8 >1 10 75 –3 0 c5 90 –1 05 20 –1 0 c8 g c2 02 -1 1 55 0 tr ia ss ic , l 23 5– 20 8 >1 20 60 –4 0 c5 95 –1 05 37 –2 1 c6 40 –6 5 13 –0 c8 g c1 03 -4 40 0 ju ra ss ic , e -m 19 0 10 0– 10 5 76 –3 8 c5 75 –8 5 14 –4 c8 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 113 of 154 www.geusbul let in.org a pp en di x 1. 4 th er m al h is to ry s ol ut io ns . a ls o av ai la bl e on lin e in s up pl em en ta ry f ile s 4 (c on tin ue d) sa m pl e nu m be r / re gi on vt el ev at io n (m a .s .l. ) st ra tig ra ph ic su bd iv is io n st ra tigr ap hi c ag e (m a) pa la eo te m p. (° c) o ns et o f co ol in g (m a) ev en t pa la eo te m p. (° c) o ns et o f co ol in g (m a) ev en t pa la eo te m p. (° c) o ns et o f co ol in g (m a) ev en t g c1 07 776 11 0 tr ia ss ic , l 23 5– 20 8 >1 15 65 –4 0 c5 10 0– 11 0 43 –1 8 c6 80 –9 5 14 –6 c8 g c1 07 777 71 7 ju ra ss ic , l 16 1– 14 6 95 –1 00 17 –3 c8 g c1 07 778 40 rh ae tia n– si ne m ur ia n 20 4– 18 9 >1 20 10 0– 48 c5 10 0– 11 0 49 –1 4 c6 60 –9 5 16 –4 c8 g c1 03 -3 66 0 ju ra ss ic , l 15 5 95 –1 05 13 0– 37 c5 85 –9 5 16 –3 c8 g c1 01 627 25 0 ju r. –c re t. 20 0– 65 90 –1 00 a c5 90 –1 00 20 –0 c8 g c1 01 628 95 tr ia ss ic 25 1– 20 0 10 0– 10 5 12 0– 15 c5 80 –1 00 13 –2 c8 6b : l iv er po ol l an d g c1 01 672 44 5 pr ec am br ./c al ed . >4 00 10 5– 11 0 30 0– 21 5 c1 85 –9 5 14 5– 25 c6 25 –8 5 50 –0 c8 g c1 01 673 55 0 pr ec am br ./c al ed . >4 00 >1 00 b ~5 5 b c5 10 0– 11 0 47 –1 9 c6 45 –9 0 22 –7 c8 g c1 07 733 vt 12 12 77 d ev on ia n, e 41 6– 39 8 >1 00 25 5– 17 0 c2 80 –1 05 20 0– 25 c6 35 –7 5 30 –0 c8 g c1 07 734 vt 12 10 58 d ev on ia n, e 41 6– 39 8 >1 00 18 0– 10 0 c2 75 –9 5 b 10 0– 10 b c6 20 –8 0 15 –0 c8 g c1 07 735 vt 12 60 0 d ev on ia n, e 41 6– 39 8 >1 00 b >1 20 b c2 90 –9 5 11 5– 25 c6 45 –8 0 30 –0 c8 g c1 07 736 vt 12 34 7 d ev on ia n, e 41 6– 39 8 >1 00 b >1 20 b c2 10 0– 11 0 75 –2 0 c6 80 –9 5 20 –0 c8 6c : b lo ss ev ill e ky st g c1 01 621 g c1 01 622 26 0 eo ce ne 49 –4 7 60 –8 5 >2 0 c6 <8 0 35 –0 c8 30 3 eo ce ne 49 –4 7 75 –1 00 >6 c6 <8 0 24 –0 c8 th er m al h is to ry s ol ut io ns d er iv ed fr om a ft a da ta b as ed o n as su m ed h ea tin g an d co ol in g ra te s of 1 °c /m a an d 10 °c /m a, r es pe ct iv el y. q uo te d ra ng es c or re sp on d to ± 95 % c on fid en ce li m its o n m ax im um /p ea k pa la eo te m pe ra tu re a nd o ns et o f c oo lin g in d is cr et e ep is od es o f h ea tin g an d co ol in g. c oo lin g ep is od es c 0– c9 a cc or di ng to t ab le 1 . a o ut ly in g co ns tr ai nt s on ti m in g. b u nc er ta in c on st ra in ts . ca le d. : c al ed on ia n. c re t.: c re ta ce ou s. e : e ar ly . j ur .: ju ra ss ic . l : l at e. m : m id dl e. v t: v er tic al tr an se ct . p os tde p. : p os tde po si tio na l. pr ec am br .: pr ec am br ia n. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 114 of 154 www.geusbul let in.org appendix 1.5 maximum palaeotemperatures from vr data in outcrop samples. also available online in supplementary file s4 sample number gc1077source number locality latitude longitude elevation (m a.s.l.) stratigraphic subdivision strat. age (ma) r0max (range) (%) eq. vr from ir (%) a n b max. palaeo temp. (°c) c cooling episode 106.1 118352 cardiocerasdal, wollaston forland 74.454 –20.140 415 bernbjerg fm (oxfordian – kimmeridgian) 161–151 0.59 (0.51– 0.68) 25 97 c4 107.1 118461 cardiocerasdal, wollaston forland 74.463 –20.165 400 bernbjerg fm (oxfordian – kimmeridgian) 161–151 0.54 (0.42– 0.68) 25 90 c4 108.1 118473 cardiocerasdal, wollaston forland 74.463 –20.165 400 bernbjerg fm (oxfordian – kimmeridgian) 161–151 0.47 (0.37– 0.56) 25 79 c4 109.1 327654 kap stosch, hold with hope 74.050 –21.700 50 ravnefjeld fm (permian) 251–299 0.39 (0.26– 0.58) 25 65 110.1 427850 stensbybjerg, gulelv hold with hope 73.900 –21.211 400 bernbjerg fm (late oxfordian – early kimmeridgian) 157–154 0.48 (0.39– 0.62) 25 80 c6 111.1 427851 stensbybjerg, gulelv hold with hope 73.900 –21.211 400 bernbjerg fm (late oxfordian – early kimmeridgian) 157–154 0.50 (0.44– 0.61) 25 83 c6 112.1 468940 sabine ø, pendulum øer 74.578 –18.987 180 paleocene 66–56 0.62 (0.51– 0.71) 25 102 c6 113.1 468950 sabine ø, pendulum øer 74.567 –18.987 400 paleocene 66–56 0.72 (0.64– 0.83) 16 118 c6 114.1 468951 sabine ø, pendulum øer 74.567 –18.987 400 paleocene 66–56 0.68 (0.60– 0.78) 25 113 c6 115.1 468954 sabine ø, pendulum øer 74.566 –18.989 445 paleocene 66–56 0.75 (0.62– 0.85) 25 122 c6 116.1 569833 brorson halvø, wollaston forland 74.598 –19.588 175 bernbjerg fm (l. jurassic) 161–146 0.64 (0.55– 0.72) 25 106 c6 117.1 469856 brorson halvø, wollaston forland 74.598 –19.588 175 cretaceous, m 145–65 0.65 (0.56– 0.77) 25 108 c6 118.1 469857 brorson halvø, wollaston forland 74.598 –19.588 175 cretaceous, m 145–65 0.62 (0.52– 0.71) 14 100 c6 119.1 470117 dronning augusta dal, wollaston forland 74.343 –19.347 304 thanetian 59-56 0.46 (0.36– 0.63) 25 78 c6 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 115 of 154 www.geusbul let in.org appendix 1.5 maximum palaeotemperatures from vr data in outcrop samples. also available online in supplementary file s4 (continued) sample number gc1077source number locality latitude longitude elevation (m a.s.l.) stratigraphic subdivision strat. age (ma) r0max (range) (%) eq. vr from ir (%) a n b max. palaeo temp. (°c) c cooling episode 120.1 470118 dronning augusta dal, wollaston forland 74.343 –19.347 304 thanetian 59–56 0.43 (0.34– 0.51) 25 72 c6 121.1 470121 dronning augusta dal, wollaston forland 74.343 –19.347 304 thanetian 59–56 0.48 (0.40– 0.55) 5 80 c6 122.1 472012 aucella bjerg, wollaston forland 74.524 –20.466 700 e. aptian – e. albian 125–112 0.51 (0.41– 0.61) 25 84 c4 123.1 472014 aucella bjerg, wollaston forland 74.524 –20.466 700 e. aptian – e. albian 125–112 0.47 (0.41– 0.54) 25 79 c4 124.1 472016 aucella bjerg, wollaston forland 74.524 –20.466 700 e. aptian – e. albian 125–112 0.43 (0.33– 0.53) 25 72 c4 125.1 492906 djævlekløften, clavering ø 74.359 –20.576 457 cretaceous, m? 145–65 4.36 (4.01– 5.01) 25 >250 c6 126.1 492907 djævlekløften, clavering ø 74.359 -20.576 457 cretaceous, m? 145–65 4.45 (3.94– 5.04) 25 >250 c6 127.1 492930 dolomitdal, clavering ø 74.378 –20.593 304 cretaceous 145-65 0,5 1 83 c6 128.1 492931 dolomitdal, clavering ø 74.378 –20.593 304 cretaceous 145-65 1.42 (0.84– 1.86) 0.54 25 90 c6 129.1 492933 dolomitdal, clavering ø 74.378 –20.593 304 cretaceous 145–65 0.54 (0.45– 0.61) 25 90 c6 130.1 493938 store koldewey 76.232 –18.579 80 cretaceous 145–65 0.55 (0.44– 0.68) 25 91 c4 131.1 493941 store koldewey 76.232 –18.579 80 cretaceous 145–65 0.55 (0.49– 0.63) 14 91 c4 132.1 493942 store koldewey 76.232 –18.579 80 cretaceous 145–65 0.54 (0.43– 0.66) 12 90 c4 133.1 493960 store koldewey 76.297 –18.634 77 cretaceous 145–65 0.43 (0.36– 0.51) 25 72 c4 134.1 493961 store koldewey 76.297 –18.634 77 cretaceous 145–65 0.51 (0.42– 0.59) 6 84 c4 http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 116 of 154 www.geusbul let in.org appendix 1.5 maximum palaeotemperatures from vr data in outcrop samples. also available online in supplementary file s4 (continued) sample number gc1077source number locality latitude longitude elevation (m a.s.l.) stratigraphic subdivision strat. age (ma) r0max (range) (%) eq. vr from ir (%) a n b max. palaeo temp. (°c) c cooling episode 135.1 493962 store koldewey 76.297 –18.634 77 cretaceous 145–65 0.58 (0.55– 0.60) 2 96 c4 136.1 495913 store koldewey 76.143 –18.542 77 cretaceous 145–65 0.51 (0.40– 0.64) 25 84 c4 137.1 495915 store koldewey 76.143 –18.542 77 cretaceous 145–65 0.51 (0.41– 0.64) 25 84 c4 138.1 495915 store koldewey 76.143 –18.542 77 cretaceous 145–65 0.51 (0.41– 0.62) 25 84 c4 139.1 516065 kap ehrenberg, payer land 74.413 –21.867 120 carboniferous 359–299 0.73 25 120 140.1 516083 kuplen, wollaston forland 74.342 –20.086 327 e. cretaceous 146–100 0.61 (0.53– 0.68) 7 100 c6 141.1 516085 kap schumacher, wollaston forland 74.645 –20.054 50 l. jurassic 161–146 0.54 (0.40– 0.67) 25 90 c6 142.1 517106 store koldewey 76.380 –18.728 54 barremian 130–125 0.49 (0.39– 0.60) 25 81 c4 143.1 517107 store koldewey 76.380 –18.728 54 barremian 130–125 0.48 (0.41– 0.56) 25 80 c4 144,1 523402 stensbybjerg, hold with hope 73.918 –21.157 770 stensbybjerg fm rødelv mb (aptian – early albian) 125–108 0.54 (0.41– 0.65) 25 90 c6 145.1 523404 stensbybjerg, hold with hope 73.918 -21.157 770 stensbybjerg fm gulelv mb (late barremian – aptian) 127–125 0.46 (0.32– 0.57) 3 78 c6 a equivalent vr derived from mean inertinite reflectance. b number of fields measured. a target of 25 fields is usually considered to charac terise an analysis of the highest quality. c all estimates of maximum palaeotemperature were determined using an assumed heating rate of 1°c/myr (sweeney & burnham 1990). cooling episodes c0–c9 according to table 1. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 117 of 154 www.geusbul let in.org c. 240 c1 95–90 c4 37–35 c6 c. 10 c8 c. 180 c2 145–140 c3 gc1016-19 gc1104-2 gc1104-3 gc1077-11 gc1104-1 gc1077-10 gc1077-9 gc1077-3 gc1077-13 gc1077-12 gc1104-22 gc1104-23 gc1077-8 gc1077-1 gc1077-5 gc1077-6 gc1077-4 gc1077-17 gc1077-15 gc1077-16 ? ? ? 320–300 c0 onset of cooling (ma): cooling episode: gc1077-104 gc1077-101 gc1077-102 gc1077-83 gc1077-103 gc1077-60 gc1077-61 time (ma)region 1 050100150200250300350400 germania land etc. store koldewey hochstetter forland etc. stratigraphic age onset of cooling in individual sample uncertain constraints onset of cooling across the region outlying value samples along a vertical transect fig a-2-region1 vt1–vt17 gc1104-20 appendix 1.6.1 timing constraints derived from afta data in individual samples in region 1. timing constraints derived from afta data in individual samples the constraints are organised into six regions (appendix 1.6.1–1.6.6) from north to south along the coast, locations in fig. 11a. horizontal coloured bars represent 95% confidence intervals on the onset of cooling, as illustrated in appendix 1.1. vertical columns represent the constraints for the onset of each of the ten regional cooling episodes defined assuming that the episodes are regionally synchronous (table 1). time-temperature constraints for each sample are listed in appendix 1.4. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 118 of 154 www.geusbul let in.org gc1016-50 gc1016-51 gc1016-53 gc1016-54 gc1016-48 gc1016-20 clavering ø gc1016-49 ? ? gc522-8 gc1016-3 wollaston forland gc522-7 gc1016-23 gc1016-52 gc1016-55 ? gc1077-22 gc1077-23 gc1077-24 gc1077-25 gc1077-26 kuhn ø gc1077-27 gc1077-28 gc1077-29 gc1077-30 gc1077-31 gc1077-32 payer land c. 240 c1 95–90 c4 37–35 c6 c. 10 c8 c. 180 c2 145–140 c3 320–300 c0 c. 55 c5 gc1077-98 gc1077-99 gc1077-72 gc1077-73 gc1077-74 gc1077-75 gc1077-100 ? ? gc1077-91 gc1077-89 gc1077-90 gc1077-92 v t1 6 v t9 ? ? gc1077-88 gc1077-84 gc1077-105 gc1077-70 gc1077-49 gc1077-47 ?? v t1 0 gc1077-80 gc1077-81 gc1077-82 gc1077-96 gc1077-95 gc1077-94 gc1077-93 gc1077-97 v t1 7 v t1 ? ? gc1077-86 gc1077-87 ? time (ma) 050100150200250300350400 region 2 c. 5 c9 20–18 c7 onset of cooling (ma): cooling episode: ? fig a-2-region2 appendix 1.6.2 timing constraints derived from afta data in individual samples in region 2. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 119 of 154 www.geusbul let in.org c. 55 c5 c. 240 c1 37–35 c6 c. 10 c8 c. 180 c2 gc1016-60 gc1016-62 gc1016-41 gc1016-57 gc1016-58 gc1016-59 145–140 c3 gc1016-63 gc1016-65 gc1016-70 gc1016-24 gc1016-25 gc1016-26 gc1016-61 gc1016-39 gc1016-68 gc1016-67 gc1016-69 gc1016-66 gc1016-38 hold with hope ? ? ? ? ? ? gc1077-40 gc1016-16 gc1016-17 gc1016-18 ? gc1016-37 gc1016-36 geog. soc. ø gc1016-35 gc1077-79 v t2 v t3 v t4 v t4 a gc1077-85 gc1077-46 gc1077-45 gc1077-44 gc1077-43 hudson land gc1077-42 gc1077-51 gc1077-50 v t5 ? ? 95–90 c4 time (ma)region 3 050100150200250300350400 ymer ø onset of cooling (ma): cooling episode: gc1104-24 gc1104-15 fig a-2-region3 appendix 1.6.3 timing constraints derived from afta data in individual samples in region 3. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 120 of 154 www.geusbul let in.org time (ma)region 4 050100150200250300350400 gc1016-4 gc1016-46 gc1016-47 gc1016-44 gc1016-45 gc1016-43 gc1016-42 gc1016-6 gc1016-7 ?? ? gc1016-15 gc522-3 gc1077-63 gc1016-8 gc1016-9 gc1077-71 gc1077-66 gc1077-62 gc1077-65 gc1077-67 gc1077-68 gc1077-69 gc1077-64 gc1077-57 gc1077-56 gc1077-58 gc1077-59 gc1077-41 gc1016-5 ? ? ? ? ? ? ? ? c. 240 c1 95–90 c4 37–35 c6 c. 10 c8 c. 180 c2 145–140 c3 320-300 c0 ? ? ? ? inland vt 8 gc1016-14 gc1016-13 ? onset of cooling (ma): cooling episode: gc1104-14 gc1104-4 gc1104-5 gc1104-6 fig a-2-region4 gc1104-21 appendix 1.6.4 timing constraints derived from afta data in individual samples in region 4. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 121 of 154 www.geusbul let in.org appendix 1.6.4 timing constraints derived from afta data in individual samples in region 4. gc522-9 gc522-1 gc522-4 gc522-11 gc522-10 traill ø schuchert dalgc103-6 gc159-1 gc159-2 gc202-1 gc202-2 gc202-3 gc202-4 gc202-5 gc202-6 gc202-7 gc202-16 gc103-2 gc103-1 gc159-3 gc159-4 wegener halvø gc522-5 gc1016-10 gc1016-11 mestersvig gc1077-18 gc1077-20 gc1077-21 gc1016-30 gc1016-29 gc1016-33 gc1016-32 gc1077-52 gc1077-54 gc1077-55 traill ø c. 240 c1 37–35 c6 c. 5 c9 c. 10 c8 c. 180 c2 c. 55 c5 20–18 c7 time (ma)region 5 050100150200250300350400 v t7 v t1 1 v t6 v t1 5 gc522-6 gc522-2 95–90 c4 onset of cooling (ma): cooling episode: gc1104-19 gc1104-16 gc1104-17 gc1104-18 gc1104-13 gc1104-12 gc1104-10 gc1104-7 gc1104-8 gc1104-9 320–300 c0 stauning alperstauning alper fig a-2-region5 appendix 1.6.5 timing constraints derived from afta data in individual samples in region 5. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 122 of 154 www.geusbul let in.org gc202-9 gc103-3 gc103-4 gc202-11 gc202-12 jameson land gc1016-27 gc1016-28 ? liverpool landgc1016-72 gc1016-73 gc1016-21 gc1016-22 gc1019-32 blosseville kyst gc1077-33 gc1077-34 gc1077-35 gc1077-36 gc1077-37 gc1077-38 gc1077-39 gc1077-76 gc1077-77 gc1077-78 ? ? ? c. 240 c1 37–35 c6 c. 10 c8 c. 180 c2 c. 55 c5 v t1 3 v t1 2 time (ma)region 6 050100150200250300350400 onset of cooling (ma): cooling episode: fig a-2-region6 appendix 1.6.6 timing constraints derived from afta data in individual samples in region 6. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 123 of 154 www.geusbul let in.org appendix 1.6.6 timing constraints derived from afta data in individual samples in region 6. appendix 1.7 thermal history interpretations for samples of the traill ø region. also available online in supplementary file s4 sample number elevation (m a.s.l.) c6 palaeo temp. (°c) c6 onset of cooling (ma) c8 palaeo temp. (°c) c8 onset of cooling (ma) c9 palaeo temp. (°c) c9 onset of cooling (ma) zone 0 gc522-6 150 90–100 45–10 20–80 16–0 gc522-2 30 90–100 125–30 60–85 30–0 gc1016-14 1 30–80 40–0 gc1016-16 1 70–90 b 100–20 b 70–80 15–0 gc1016-17 1 85–95 70–15 50–75 12–0 gc1016-18 1 95–105 55–15 70–90 15–5 zone 1 gc1016-37 850 65–85 35–0 gc1016-36 370 40–85 20–0 gc1016-35 300 80–85 32–2 gc1077-50 255 100–105 70–45 a 90–95 15–5 gc1077-42 160 >105 30–5 a gc522-9 140 >110 55–20 90–105 15–5 gc1077-51 103 100–110 40–15 80–95 10–3 zone 2 gc1016-30 1075 105–110 45–30 80–90 20–10 gc1104-19 1069 90–105 43–19 45–75 10–0 gc1016-29 730 >100 55–20 80–100 12–3 gc1104-17 570 >105 55–25 <105 30–0 gc1104-16 570 90–115 43–10 30–80 12–0 gc1104-15 570 45–75 30–0 gc1016-33 460 >130 42–22 90–105 13–2 gc1016-32 323 >125 38–25 80–95 12–4 gc522-10 230 >120 55–20 100–110 27–5 40–90 7–0 zone 3 gc1077-52 1018 >105 65–20 <105b 20–0 b gc1077-54 689 65–100 25–0 gc522-1 500 >125 45–25 55–80 22–7 a gc1104-18 471 >120 40–25 105–110 14–4 30–70 17–0 gc522-11 300 >110 50–15 105–115 28–10 70–90 5–0 gc522-4 140 >125 ~35 >110 22–5 35–90 13–0 gc1077-55 128 55–95 8–0 zones 0–3 are defined in figure 32. cooling episodes c6, c8 and c9 according to table 1. a outlying constraints on timing. b uncertain constraints. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 124 of 154 www.geusbul let in.org appendix 2 results from vertical transects individual vertical transects samples were collected along 17 near-vertical transects (vt1 to vt17; locations in appendix 2.1). geological constraints on the burial and exhumation history for the samples in each vertical transect are summarised in appendix 2.2. for each transect we omitted samples from too wide a horizontal distance to guard against lateral variation in the magnitude of palaeothermal effects, which might arise due to variation in heat flow, differential uplift (tilting) or offsets across major faults. comparison of results from adjacent profiles shows that this approach is justified. the palaeothermal constraints for the samples along each transect are shown in appendix 2.3. for a number of these transects, sufficient constraints are available to justify quantitative determination of palaeogeothermal gradients and removed sections (appendix 2.4). palaeotemperature constraints for the mesozoic and end-eocene episodes from the various transects are compared in appendix 2.5. traill ø 9 1 23 4 14 10 13 7 8 11 16 17 12 5 15 6 fig. b-1 vt1: clavering ø (w) vt2: hold with hope (nw) vt3: hudson land (n) vt4,4a: hudson land (s) vt5: geographical society ø vt6: traill ø (s) vt8: milne land vt9: kuhn ø vt10: payer land vt11: stauning alper vt12: liverpool land vt13: jameson land vt14: giesecke bjerge vt15: traill ø (e) vt16: dombjerg vt17: clavering ø (ne) 30°w 20°w 30°w 20°w 76 °n 74 °n 72 °n 76 °n 74 °n 72 °n 70 °n 100 km appendix 2.1 location of vertical transects vt1 to vt14. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 125 of 154 www.geusbul let in.org appendix 2.2 geological constraints on the burial and exhumation history of vertical transects provided in appendix 2.3. locations in appendix 2.1 vt area palaeozoic mesozoic cenozoic vt1 clavering ø (w) upper carboniferous sediments a (on basement) palaeogene basalts (summits) vt2 hold with hope (nw) upper permian sediments a (on basement) triassic sediments a palaeogene basalts (summits) vt3 hudson land (n) upper carboniferous sediments a (on basement) vt4 hudson land (e) vt5 geographical society ø triassic, lower cretaceous sediments a vt6 traill ø (sw) jurassic – lower cretaceous sediments a vt7 schuchert dal carboniferous–permian sediments a triassic sediments a vt8 milne land middle jurassic sediments a (on basement) palaeogene basalts (summits) vt9 kuhn ø middle jurassic sediments (on basement) palaeogene basalts (summits) vt10 payer land vt11 stauning alper vt12 liverpool land sub-permian peneplain along western liverpool land vt13 jameson land jurassic sediments a vt14 giesecke bjerge upper permian sediments a (on basement) triassic – lower cretaceous a sediments palaeogene basalts (summits) vt15 traill ø (e) palaeogene intrusives a vt16 dombjerg lower cretaceous sediments a (on basement) palaeogene basalts (summits) vt17 clavering ø (ne) upper permian carbonates (on basement in the summits) triassic – lower cretaceous sediments (in nearby rift basin) palaeogene basalts (summits) a samples included in the vertical transect (other than basement samples). http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 126 of 154 www.geusbul let in.org vt1: clavering ø (w). seven samples of the carboniferous traill ø group were collected over an elevation difference of c. 800 m from the south-west corner of clavering ø. the highest local summit reaches 1123 m a.s.l. where palaeogene basalts are present here and at other summits in this region. afta data in samples from vt1 define three palaeothermal episodes: middle triassic (c1), earliest cretaceous (c3) and late miocene (c8). palaeotemperatures characterising the earliest cretaceous and late miocene palaeothermal episodes define linear profiles, broadly consistent with palaeogeothermal gradients of c. 25°c/km. the variation of middle triassic palaeotemperatures with elevation is not clear because all but the highest elevation sample were totally annealed in this episode. the linear variation of palaeotemperatures with elevation for the earliest cretaceous and late miocene episodes suggests that these palaeotemperatures reflect additional depth of burial, and that cooling was due predominantly to exhumation. no such insights are possible into the origin of palaeotemperatures characterising the middle triassic episode, but a similar conclusion seems likely. the late miocene palaeotemperature profile intersects the local summit level at c. 40°c, which is higher than any reasonable value of surface temperature, implying that a significant amount of cover was present at the onset of this cooling episode. since palaeogene basalts are present in the summits, the additional burial required to explain the late miocene palaeotemperatures must have been deposited after the preserved basalts were laid down. this most likely represented younger basalts together with eocene–miocene sedimentary units. similarly, the earliest cretaceous palaeotemperatures intersect the summit level at a value of c. 70°c, implying that a considerable thickness of post-carboniferous succession has been eroded on the unconformity below the palaeogene basalts. the middle triassic palaeotemperatures intersect the summit level at 100–110°c. since all samples along vt1 are of upper carboniferous sediment, the heating to the middle triassic palaeotemperature was due to burial below upper carboniferous to middle triassic sediments. note that the earliest cretaceous and the late miocene events are identified in all samples. this implies that the late miocene palaeotemperatures are likely to be highly reliable. vt2: kap stosch (hold with hope). four samples of permian and triassic sedimentary rocks were collected over an elevation difference of c. 320 m at the northwest tip of hold with hope, north-west of the spath plateau, comprised of palaeogene basalts. the local summit reaches 1211 m a.s.l. where thin palaeogene basalts are present. just to the south-east, the permian– triassic succession is faulted against a huge thickness of palaeogene basalts (upton et al. 1980). afta data in samples from vt2 define two palaeothermal episodes: earliest cretaceous (c3) and late miocene (c8). the range of elevations over which palaeotemperatures have been defined is not sufficiently large to define the variation of palaeotemperatures with elevation for each episode. however, palaeotemperatures in the mid-cretaceous and late miocene episodes are similar to those at the same elevations from vt1 (western clavering ø). it seems likely that similar mechanisms of heating and cooling apply at both locations. given the narrow interval over which samples have been analysed at this location, quantitative determination of palaeogeothermal gradients and removed section is unlikely to provide tight constraints from this transect. however, these data do provide useful insights into the regional framework when compared with other transects. note that the mid-cretaceous and the late miocene episodes are identified in all samples, implying that the late miocene palaeotemperatures are highly reliable. vt3: hudson land (n). four samples of the carboniferous traill ø group were collected over an elevation difference of c. 800 m from north-east hudson land. the highest local summit reaches c. 1100 m a.s.l. and no palaeogene basalts are present. however, basalts are present at low elevations to the south-east in faulted contact with older sedimentary units. afta data in samples from vt3 define four palaeothermal episodes: early jurassic (c2), earliest cretaceous (c3), end-eocene (c6) and late miocene (c8). not all events are recognised in every sample and data from these samples lack the degree of consistency seen elsewhere, for example vt1 (western clavering ø). this is probably due to difficulties in resolving palaeotemperatures in multiple palaeothermal episodes, particularly when they are closely spaced in time and magnitude. while the earlier of these episodes is identified in samples gc1016-25 and -70, the more recent episode is recognised in samples gc1016-24 and -26, in which c3 palaeotemperatures are also identified. the early jurassic palaeotemperatures could be explained in terms of a linear palaeotemperature profile characterised by an extremely high palaeogeothermal gradient in excess of 100°c/km. this seems unlikely, given the lack of supporting evidence from similar locations. we interpret the rapid increase of palaeotemperature in this episode between samples gc1016-70 and -24 as more likely due to major tectonic offset across a major fault, which has not yet been identified. note that the end-eocene and the late miocene events are identified in the lowermost sample, and thus the late miocene palaeotemperature range for this sample is highly reliable. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 127 of 154 www.geusbul let in.org vt4, 4a: hudson land (e). six samples of precambrian basement rocks were collected over an elevation range of c. 1100 m in north-east hudson land, south of vt3. an additional three samples (two of precambrian basement and one devonian sedimentary rock, which failed to yield apatite) were collected farther to the south-west. these samples are referred to as vt4a. the highest local summit reaches c. 1500 m a.s.l., and no palaeogene basalts are present, although basalts are present at low elevations to the south-east in faulted contact with precambrian rocks (and to the north on clavering ø). afta data in samples from vt4 define four palaeothermal episodes: middle triassic (c1), earliest cretaceous (c3), end-eocene (c6) and late miocene (c8). linear variation of palaeotemperatures with elevation for the end-eocene and late miocene episodes suggests that palaeotemperatures in both episodes reflect additional depth of burial. an additional component of elevated heat flow during the end-eocene episode is also possible, since these locations are close to the late eocene igneous centre at myggbukta. palaeotemperatures in the end-eocene (c6) episode appear to define a distinct hot spot around this centre (map c6 in fig. 14). the constraints for episodes c6 and c8 are consistent with a palaeogeothermal gradient 30°c/km and 25°c/ km, respectively. pre-palaeogene episodes are identified in these transects only in isolated samples, and in all cases these samples were totally annealed prior to the onset of cooling in the respective episodes. therefore, no information is available on the variation of palaeotemperature with elevation for these episodes, or on the relevant mechanisms of heating and cooling. similar to vt1, the late miocene palaeotemperature profile for vt4 intersects the local summit level at c. 40°c. this is higher than any reasonable value of surface temperature, implying that a significant amount of cover was present at the onset of this cooling episode. similar comments apply to the end-eocene episode, for which the summit intercept is closer to 70°c. since no palaeogene basalts are present on the summit, the age of the additional section that is required to explain the late miocene and end-eocene palaeotemperatures remains uncertain at this location. end-eocene palaeotemperatures from vt4a (samples gc1016-63, -65) are consistently lower than in samples from vt4 (vt4a is located about 10 km south of vt4). this offset suggests either a difference in basal heat flow or in the amount of post-end-eocene exhumation between these locations. given the presence of downfaulted palaeogene basalts not far to the south, differential exhumation across a fault between the two locations appears likely. note that the end-eocene and the late miocene events are identified in all samples. this condition and the high late miocene palaeotemperatures from which these samples have cooled, implies that the late miocene palaeotemperatures are highly reliable. vt5: geographical society ø. two samples of lower cretaceous and one of triassic sedimentary rocks were collected over an elevation difference of c. 550 m from the centre of geographical society ø. the highest local summit reaches c. 1060 m a.s.l. and no palaeogene basalts are recognised within the vicinity. afta data in samples from vt5 define just two palaeothermal episodes: mid-cretaceous (c4) and late miocene (c8). palaeotemperatures characterising the mid-cretaceous episode are identified only in the single sample of triassic age, so no information is available on the form of the profile characterising this episode. late miocene palaeotemperatures are consistent with a linear profile defining a palaeogeothermal gradient of c. 25°c/ km, although the quite broad palaeotemperature constraints allow a range of interpretations. vt6: traill ø (sw). five samples were analysed from jurassic and cretaceous sedimentary units on traill ø, over an elevation difference of c. 800 m. the highest local summit reaches c. 1400 m a.s.l. and no palaeogene basalts are present in the immediate vicinity, although they are present on summits to the north, at lower elevations. late eocene syenites are also present at the eastern extreme of the island, forming kap simpson and kap parry. afta data in samples from vt6 define three palaeothermal episodes: end-eocene (c6), late miocene (c8) and early pliocene (c9; see also fig. 32). palaeotemperatures characterising the end-eocene and late miocene episodes define linear profiles, consistent with palaeogeothermal gradients around c. 40°c/ km and 25°c/km, respectively. the early pliocene episode is only recognised in a single sample at low elevation, presumably because palaeotemperatures in higher elevation samples were too low to be resolved with confidence. in sample gc1016-32, the late miocene palaeotemperature is lower than the trend defined for this episode by data in other samples. this may reflect the difficulty in resolving the late miocene and early pliocene episodes in a single sample. it seems likely that the solution for this sample represents the unresolved effects of these two episodes. a fault offset could provide an alternative explanation. linear variation of palaeotemperatures with elevation for the end-eocene and late miocene episodes suggests that these palaeotemperatures reflect additional depth of burial. the apparent decrease in palaeogeothermal gradient from end-eocene to late miocene suggests a significant decrease in basal heat flow over that period. thus, cooling in these episodes is interpreted as due to a combination of exhumation and a decrease in http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 128 of 154 www.geusbul let in.org basal heat flow. no such insights are possible into the origin of palaeotemperatures characterising the early pliocene episode, but it seems likely that this episode represents the final stages of exhumation. the end-eocene palaeotemperature profile intersects the local summit level at c. 90°c. the corresponding value for the late miocene profile is c. 70°c. these values imply that a significant amount of cover was present at the onset of these cooling episodes. since no palaeogene basalts are present on the local summit, the additional burial required to explain the end-eocene palaeotemperatures probably consisted of more cretaceous sedimentary units as well as basalts and eocene sedimentary units. clearly, some of the additional section remained until the onset of late miocene cooling, although some degree of re-burial between the two cooling episodes is possible. note that the end-eocene and the late miocene events are identified in all samples. together with the high late miocene palaeotemperatures, from which these samples have cooled, this implies that the late miocene palaeotemperatures are highly reliable. vt7: schuchert dal. eleven samples of carboniferous to jurassic sedimentary units were analysed over an elevation difference of more than 1000 m from schuchert dal, western jameson land. the highest local summits in the vicinity of the samples reach c. 2000 m a.s.l. but peaks up to 2.5 km are present to the west in stauning alper. palaeogene basalts are not present in the vicinity of these samples, although the igneous centre of werner bjerge is not far away to the north-east. precise locations for the samples were not available, so the co-ordinates adopted for these samples in this study were obtained by digitising maps from hansen et al. (2001), from which elevations were derived. unfortunately, those maps are not sufficiently detailed to identify precise co-ordinates, and elevations obtained by this process are different to those published by hansen et al. (2001), which in turn are different from those in the original geotrack reports (see green 2009). the results do not define consistent palaeotemperature profiles. it is thus uncertain whether this lack of consistency reflects real complexity in the nature of the palaeothermal processes, or uncertainty in sample locations. afta data in samples from vt7 define three palaeothermal episodes: early eocene (c5), early miocene (c7) and early pliocene (c9). palaeotemperatures characterising all three palaeothermal episodes define highly nonlinear profiles, suggesting that a simple interpretation of all data as representing a single coherent rock mass that has undergone conductive heating and cooling is not tenable. the profile characterising the early eocene (c5) episode in particular shows erratic variation through the section. we interpret this episode as reflecting the effects of contact heating and hydrothermal effects associated with palaeogene igneous activity, which explains this erratic variation through the section. the early miocene (c7) cooling episode occurs slightly later than intrusive activity recognised at werner bjerge dated at 26.8 ± 0.1 ma age (brooks et al. 2004). however, the close proximity between the sample reflecting early miocene cooling and the intrusive centre (map c7 in fig. 14), suggests a direct link between the two. this may explain the non-linear variation of palaeotemperatures with elevation in these samples, where distance to the intrusive centre as well as possible late-stage hydrothermal effects will also affect the magnitude and timing of this episode. the early pliocene cooling episode is only identified in two samples, but even these do not show a progressive increase with depth through the section. alternatively, it is quite possible that the elevations are incorrect. other explanations involving differential exhumation or localised heating and cooling are possible, but unlikely. another problem in deriving reliable thermal history constraints from afta data in the vicinity of vt7 is the convergence of the effects of a number of palaeothermal episodes in this region, particularly the early eocene and early miocene episodes. both of these episodes appear to involve a significant enhancement of basal heat flow associated with intrusive activity, as well as the early eocene episode, which we also attribute largely to localised igneous activity. the interplay between these competing processes, involving broadly similar palaeotemperatures in the vicinity of vt7, makes it quite likely that the effects of one episode may have been attributed to another, or that individual events identified from afta data may in fact represent unresolved effects of both. the consistent definition of the early miocene episode in samples from vt7 and adjacent locations, and the resulting highly consistent association between this episode and the werner bjerge intrusion (fig. 5), suggests that the sample locations are broadly correct. however, given the uncertainties, we have not attempted any quantitative analysis of palaeotemperature profiles from this transect. vt8: milne land. three samples of ordovician granite-gneiss and one sample from a jurassic outlier were collected over an elevation difference of c. 750 m from the south coast of milne land. the highest local summit reaches 1720 m a.s.l. and palaeogene basalts, up to 700 m thick, are present. four palaeothermal episodes are recognised in the afta data in samples from vt8: middle triassic (c1), early jurassic (c2), mid-cretaceous (c4) and late miocene (c8). palaeotemperatures characterising the middle triassic and early jurassic episodes are only definitively http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 129 of 154 www.geusbul let in.org recognised in single samples of the ordovician granite-gneiss, so the variation of palaeotemperature with elevation for these episodes cannot be defined. similarly, the mid-cretaceous episode is only identified in one sample (gc1042-43), although this episode is identified from the afta data from other adjacent locations. the lack of evidence for this episode in other samples is most likely due to the influence of competing episodes. the early jurassic palaeotemperature of c. 100°c for sample gc1016-45, which is located near a middle jurassic outlier, indicates that significant exhumation took place between the onset of episode c2 and the deposition of sediments in the middle jurassic. palaeotemperatures characterising the late miocene episode in samples from this transect define a linear profile, broadly consistent with a palaeogeothermal gradient of c. 25°c/km. the form of this profile suggests that these palaeotemperatures reflect additional depth of burial, and that cooling was due predominantly to exhumation. the late miocene palaeotemperature profile intersects the local summit level at c. 40°c, which is higher than any reasonable value of surface temperature. however, only sample gc1016-43 defines the mid-cretaceous event. regional analysis of the variation of palaeotemperatures in the late miocene episode relative to elevation (fig. 22) reveals that palaeotemperatures for samples -42 and -44 represent the combined effects of unresolved cooling episodes. vt9: kuhn ø. five samples of early devonian granite were collected over an elevation difference of c. 665 m from kuhn ø. the highest local summit lies to the south at an elevation of 1300 m a.s.l. where palaeogene basalts are present. afta data in samples from vt9 define three palaeothermal episodes: middle triassic (c1), mid-cretaceous (c4) and late miocene (c8). palaeotemperatures characterising the mid-cretaceous and late miocene palaeothermal episodes are broadly consistent with palaeogeothermal gradients of c. 25°c/km. the form of the profile characterising the middle triassic episode is not as well defined because all samples were totally annealed in this episode and provide only minimum values. the linear variation of palaeotemperatures with elevation for the earliest cretaceous and late miocene episodes suggests that these palaeotemperatures reflect additional depth of burial, and that cooling was due predominantly to exhumation. no such insights are possible into the origin of palaeotemperatures characterising the middle triassic episode, but a similar conclusion seems likely. the late miocene palaeotemperature profile intersects the local summit level between 30 and 40°c, which is higher than any reasonable value of late miocene surface temperature. in contrast to vt8, however, all samples from this profile define the mid-cretaceous (c4) episode, and it is reasonable to assume that the c8 palaeotemperatures are reliable. this implies that a significant amount of cover was present at the onset of this cooling episode. since palaeogene basalts are present on the summit, the additional section required to explain the late miocene palaeotemperatures must have been deposited after the preserved basalts were laid down, and most likely represented younger basalts together with eocene–miocene sedimentary units. similarly, the mid-cretaceous palaeotemperatures intersect the summit level at c. 60°c. middle triassic values are higher, implying that a considerable thickness of granite and possible post-devonian cover has been eroded at the unconformity below the palaeogene basalts in these two episodes. vt9 is sampled on a basement block covered by middle jurassic and younger sediments on its western flank, not far from the summit sample. thus, a significant column of caledonian basement to lower triassic sediments was removed between the middle triassic and the deposition of middle jurassic sands. since middle jurassic sediments are not far from the summits of kuhn ø, the mid-cretaceous palaeotemperature of c. 60°c at summit level corresponds to a heating of 40°c below a middle jurassic to lowermost upper cretaceous cover around 95 ma (surface temperature 20°c). the thickness of this cover was about 1.5 km for a gradient of 25°c/km, and it was removed prior to eruption of basalts. vt10: payer land. six samples of palaeozoic basement rocks were collected over an elevation difference of more than 1200 m on payer land. the highest local summit is close to the highest sample at an elevation of c. 1300 m a.s.l. no palaeogene basalts are present in the immediate vicinity of this transect. afta data in samples from vt10 define five palaeothermal episodes: early jurassic (c2), earliest cretaceous (c3), mid-cretaceous (c4), late miocene (c8) and early pliocene (c9). the two cretaceous episodes and the early pliocene episode are represented only in single samples in each case, and do not define the form of the corresponding palaeotemperature profiles. palaeotemperatures characterising the late miocene palaeothermal episode define a linear profile consistent with palaeogeothermal gradients of c. 25°c/km. the variation of palaeotemperature with elevation for the early jurassic episode is not defined because all samples were totally annealed in this episode and provide only minimum values. note that the single constraint attributed to the mid-cretaceous episode in sample gc1077-32, close to sea level, plots around the profile of late miocene palaeotemperatures at higher elevations. this seems to be due to the competing effects of multiple unresolved http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 130 of 154 www.geusbul let in.org cooling episodes in the data and may indicate that the c8 palaeotemperatures in sample -32 and above are overestimated due to unresolved effects of episode c4. in addition, the pliocene episode is identified only in this sample. attribution to the pliocene episode is largely based on the mismatch between this and the projected late miocene values. linear variation of palaeotemperatures with elevation for the late miocene episode suggests that these palaeotemperatures reflect additional depth of burial, and that cooling was due predominantly to exhumation. no such insights are possible into the origin of palaeotemperatures characterising the early jurassic episode, but a similar conclusion seems likely. the late miocene palaeotemperature profile intersects the local summit level at c. 45°c. mid-cretaceous palaeotemperatures at this location are likely to have been of similar magnitude. earliest cretaceous and jurassic palaeotemperature profiles intersect the summit level at values of c. 70°c and c. 100°c, respectively, implying erosion of a considerable thickness of post-palaeozoic cover in these two episodes. vt11: stauning alper. four samples of precambrian basement rocks were collected over an elevation difference of over 1750 m from stauning alper. one failed to yield any apatite suitable for analysis (gc1077-19). the highest local summits rise to c. 2100 m a.s.l. and higher. no palaeogene basalts are present in the region. afta data in samples from vt11 define three palaeothermal episodes: early jurassic (c2), mid-cretaceous (c4) and late miocene (c8). the mid-cretaceous episode is represented only in a single sample. the late miocene episode is identified in only two samples, which do not define a linear palaeotemperature profile (fig. 22). the profile characterising the early jurassic episode is not well defined since all samples were totally annealed in this episode and provide only minimum values of palaeotemperature. for this reason, results from this transect provide only very broad constraints on the palaeogeothermal gradients and removed section. however, these data do provide useful insights into the regional framework when compared with results from other transects. vt12: liverpool land. four samples of caledonian basement rocks were collected over an elevation difference of over 930 m from liverpool land. the highest local summit reaches c. 1300 m a.s.l., and no palaeogene basalts are present in the region. afta data in samples from vt12 define three palaeothermal episodes: early jurassic (c2), end-eocene (c6) and late miocene (c8). palaeotemperatures characterising the end-eocene and late miocene palaeothermal episodes define linear profiles, consistent with palaeogeothermal gradients of c. 25°c/km, although a wide range would be allowed. the variation of early jurassic palaeotemperatures with elevation is not well defined because all samples were totally annealed in this episode and provide only minimum values. the linear variation of palaeotemperatures with elevation for the end-eocene and late miocene episodes suggests that these palaeotemperatures reflect additional depths of burial, and that cooling was due predominantly to exhumation. no such insights are possible into the origin of palaeotemperatures characterising the early jurassic episode, but a similar conclusion seems likely. the late miocene palaeotemperature profile intersects the local summit level at c. 60°c. this is higher than any reasonable value of late miocene surface temperature, implying that a significant amount of cover was present at the onset of this cooling episode. similar comments apply to the end-eocene episode. the sub-permian peneplain extends across the western flanks of liverpool land (see chapter 2), but it is heavily dissected towards the atlantic margin. this implies that the summits around vt12 are located near the sub-permian peneplain and that the samples at highest elevation were near the surface in mid-permian times. consequently, the early jurassic palaeotemperatures in excess of 100°c for these samples represent heating below a cover of upper permian to lower jurassic sediments. note that both the end-eocene and the late miocene events are identified in all samples, indicating that late miocene palaeotemperatures are highly reliable. vt13: jameson land. three samples of jurassic sedimentary rocks were collected over an elevation difference of c. 750 m from jameson land. the highest peak in the region lies at an altitude of c. 1245 m a.s.l. no palaeogene basalts are present in the region. afta data in samples from vt13 define three palaeothermal episodes: early eocene (c5), end-eocene (c6) and late miocene (c8). each episode is represented in only two of the three samples, but no episode is constrained in all three. given the narrow interval over which samples have been analysed at this location, quantitative determination of palaeogeothermal gradients and removed section is unlikely to provide tight constraints. however, comparison with results from other transects (e.g. vt11 and vt12) support the attributions shown here. note the good definition of the late miocene episode in sample gc1077-39 for which the end-eocene (c6) episode is also identified. in contrast, episode c6 is not identified in sample -38, and this probably explains the broad interval of palaeotemperatures allowed for episode c8 in this sample. vt14: giesecke bjerge. two samples of basement rocks and two from overlying permian and triassic http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 131 of 154 www.geusbul let in.org sedimentary rocks were collected over an elevation difference of c. 540 m from giesecke bjerge, to the southeast of hudson land and south-west of hold with hope. the highest summit in the region is c. 1 km a.s.l., and the mountains are capped by extensive palaeogene basalts. the sedimentary section contains numerous intrusions. afta data in samples from vt14 define two palaeothermal episodes: end-eocene (c6) and late miocene (c8). the late miocene episode is represented in three samples over a very restricted vertical range, while all samples were totally annealed in the end-eocene episode and provide only minimum palaeotemperatures. these results therefore do not define the form of the corresponding palaeotemperature profiles in sufficient detail to allow quantitative analysis. however, these data do provide useful insights into the regional framework when compared with results from other transects. both the end-eocene and the late miocene events are identified in three samples, and this indicates that late miocene palaeotemperatures are highly reliable. vt15: traill ø (e). four samples of late eocene syenite were collected over an elevation difference of c. 900 m from japetus bjerg, eastern traill ø. the highest summit in the region lies at an altitude of 1042 m a.s.l. the transect and surrounding region consists entirely of late eocene intrusive rocks. one sample failed to yield any apatite suitable for analysis. afta data in the samples from vt15 define three palaeothermal episodes: end-eocene (c6), late miocene (c8) and early pliocene (c9). the late miocene palaeotemperatures appear to define a linear profile and can be described by a palaeogeothermal gradient around 40°c/km, although constraints are available in only three samples. a wide range of palaeogradients would be allowed, given that the highest elevation sample provides only a maximum limit to the allowed palaeotemperature in this episode and the lowest elevation sample provides only a minimum limit. therefore, it is difficult to reach firm conclusions regarding the nature of this episode from these data alone. in addition, the end-eocene episode is recognised only in a single sample, and thus provides only a minimum limit to the palaeotemperature in this episode. thus, these results do not define the form of the corresponding palaeotemperature profiles in sufficient detail to allow quantitative analysis. the data do, however, provide useful insights into the regional framework when compared with results from other transects and with results from other samples in the traill ø area (fig. 32). the end-eocene and the late miocene events are identified in all samples. this and the high palaeotemperatures in the late miocene imply that these palaeotemperatures are highly reliable. vt16: dombjerg. three samples of caledonian crystalline basement and one sample of overlying lower cretaceous conglomerate were collected over an elevation difference of over 1200 m from dombjerg, north of clavering ø. the highest local summit reaches c. 1500 m a.s.l. and is capped by palaeogene basalt. afta data in samples from vt16 define five palaeothermal episodes: early jurassic (c2), earliest cretaceous (c3), mid-cretaceous (c4), early eocene (c5) and late miocene (c8). afta data in sample gc1077-91 provide only tentative evidence for the early jurassic episode. the earliest cretaceous episode is only identified definitively in one sample and tentatively in another. therefore, no information is available on the form of the palaeotemperature profiles characterising these episodes. the early eocene episode is identified in two samples, but a major difference in c5 palaeotemperature is evident between the two samples, which are separated by only a small difference in elevation. this episode is interpreted as representing contact and/or hydrothermal heating associated with eocene intrusives in the region. late miocene palaeotemperatures define a linear profile, consistent with a palaeogeothermal gradient of c. 25°c/km. although only two samples define the mid-cretaceous episode at this location, the palaeotemperatures characterising this episode are consistent with a similar profile. linear variation of palaeotemperatures with elevation characterising the late miocene and mid-cretaceous episodes suggests that these palaeotemperatures reflect additional depth of burial, and that cooling was due predominantly to exhumation. while no such insights into the origin of palaeotemperatures characterising the earliest cretaceous and early jurassic episodes are possible, a similar conclusion seems likely for these. vt16 is sampled along the flank of an early cretaceous fault block that was covered by lower cretaceous sediments (gc1077-89). the mid-cretaceous palaeotemperature (75–90°c) for the summit sample is therefore due to burial below a cover of lower cretaceous to turonian sediments. the late miocene palaeotemperature profile intersects the local summit level at a value around 30 to 40°c. this is higher than any reasonable value of late miocene surface temperature, implying that a significant amount of cover of basalts and eocene¬–miocene sediments were present at the onset of this cooling episode. vt17: clavering ø (ne). four samples of basement rocks and one sample of overlying paleocene sandstones were collected over an elevation difference of c. 800 m from north-east clavering ø. the highest local summit reaches 1370 m a.s.l. palaeogene basalts are present here and at other summits at similar elevations in this region. a sample of palaeogene sediments (gc1077-96) http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 132 of 154 www.geusbul let in.org preserved below the basalts, is included in this transect. afta data in samples from vt17 define three palaeothermal episodes: middle triassic (c1), early jurassic (c2) and late miocene (c8). palaeotemperatures characterising the late miocene palaeothermal episode define a linear profile, broadly consistent with palaeogeothermal gradients of c. 25°c/ km. the early jurassic episode is recognised in only two samples from very similar elevations, so the profile characterising this episode is not well defined. similarly, the form of the profile characterising the middle triassic episode is not well defined because all samples were totally annealed prior to the onset of cooling. but results for each of these episodes would be consistent with profiles similar to that defined for the late miocene. the linear variation of palaeotemperatures with elevation for the late miocene episode suggests that these palaeotemperatures reflect additional depth of burial, and that cooling was due predominantly to exhumation. no such insights are possible into the origin of palaeotemperatures characterising the middle triassic episode, but a similar conclusion seems likely. the late miocene palaeotemperature profile intersects the local summit level at c. 60°c. this is higher than any reasonable value of surface temperature, implying that a significant amount of cover was present at the onset of this cooling episode. since palaeogene basalts are present on nearby summits, the additional section required to explain the late miocene palaeotemperatures must have been deposited after the preserved basalts were laid down. the section most likely consisted of younger basalts together with eocene–miocene sedimentary units. similarly, the early jurassic palaeotemperature profile intersects the summit level at c. 80°c while middle triassic values are higher still (assuming a palaeogeothermal gradient of 25°c/km). this implies that a considerable thickness of post-carboniferous succession has been eroded at the unconformity below the palaeogene basalts in these two episodes. the mid-cretaceous episode is, however, not resolved in any of the vt17 samples, and it is thus likely that the late miocene palaeotemperatures are affected by unresolved effects of multiple episodes. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 133 of 154 www.geusbul let in.org 0.5 1.0 0.0 1.5 timing (ma): episode: palaeotemperature (°c) 20 40 60 80 100 120 140 160 1800 2.0 el ev at io n (k m ) 0.5 1.0 0.0 1.5 timing (ma): episode: palaeotemperature (°c) 20 40 60 80 100 120 140 160 1800 2.0 el ev at io n (k m ) 0.5 1.0 0.0 1.5 timing (ma): episode: palaeotemperature (°c) 20 40 60 80 100 120 140 160 1800 2.0 el ev at io n (k m ) 0.5 1.0 0.0 1.5 timing (ma): episode: palaeotemperature (°c) 20 40 60 80 100 120 140 160 1800 2.0 el ev at io n (k m ) 0.5 1.0 0.0 1.5 timing (ma): episode: palaeotemperature (°c) 20 40 60 80 100 120 140 160 1800 2.0 el ev at io n (k m ) 0.5 1.0 0.0 1.5 timing (ma): episode: palaeotemperature (°c) 20 40 60 80 100 120 140 160 1800 2.0 el ev at io n (k m ) u pp er c ar bo ni fe ro us u pp er c ar bo ni fe ro us pe rm o– tr ia ss ic ba se m en t fig b-2_jbl vt1: clavering ø (w) c. 10 145–140 c. 10 145–140 gc1016-49 gc1016-50 gc1016-51 gc1016-53 gc1016-52 gc1016-54 gc1016-55 c. 240 c8 c3 c8 c3c1 25°c/km gc1016-68 gc1016-67 gc1016-69 gc1016-66 25°c/km vt2: kap stosch (hold with hope) vt3: hudson land (n) c. 10 145–140 gc1016-70 gc1016-24 gc1016-26 gc1016-25 c. 180 25°c/km 37–35 c8 c3 c2c6 vt4,4a: hudson land (e) c. 10 37–35 145– 140 gc1016-57 gc1016-58 gc1016-62 gc1016-61 c. 240 c8 c6 c3 c1 gc1016-60 gc1016-59 gc1016-63 gc1016-65 25°c/km 30°c /km 25°c /km 25°c /km vt5: geographical society ø c. 10 95–90 c8 c4 gc1016-37 gc1016-36 gc1016-35 25°c/km faulttr ia ss ic lo w er c re ta ce ou s ju ra ss ic – l ow er c re ta ce ou s vt6: traill ø (sw) gc1016-30 gc1016-32 gc1016-29 gc1016-33 c. 10 37–35 c. 40°c/km gc522-10 25°c/km 25°c/km c. 5 c8 c6c9 appendix 2.3 palaeotemperature constraints derived from afta data from vertical transects 1 to 17 plotted against elevation (a.s.l.). horizontal lines indicate range of palaeotemperatures estimated from afta data. arrows indicate lack of upper or lower limit for the estimated range. see appendix 2.2 for a summary of the geological constraints on the burial and exhumation history at the locations of these transects. locations in appendix 2.1. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 134 of 154 www.geusbul let in.org 0.5 1.0 0.0 1.5 timing (ma): episode: palaeotemperature (°c) 20 40 60 80 100 120 140 160 1800 2.0 el ev at io n (k m ) 0.5 1.0 0.0 1.5 timing (ma): episode: palaeotemperature (°c) 20 40 60 80 100 120 140 160 1800 2.0 el ev at io n (k m ) 0.5 1.0 0.0 1.5 timing (ma): episode: palaeotemperature (°c) 20 40 60 80 100 120 140 160 1800 2.0 el ev at io n (k m ) 0.5 1.0 0.0 1.5 timing (ma): episode: palaeotemperature (°c) 20 40 60 80 100 120 140 160 1800 2.0 el ev at io n (k m ) 0.5 1.0 0.0 1.5 timing (ma): episode: palaeotemperature (°c) 20 40 60 80 100 120 140 160 1800 2.0 el ev at io n (k m ) 0.5 1.0 0.0 1.5 timing (ma): episode: fig b2_continued1_jbl palaeotemperature (°c) 20 40 60 80 100 120 140 160 1800 2.0 el ev at io n (k m ) ba se m en t u pp er p al ae oz oi c – tr ia ss ic ba se m en t ba se m en t ba se m en t vt7: schuchert dal c. 5 gc202-6 c. 5520–18 c9 c5c7 gc103-6 gc202-7 gc202-3 gc202-16 gc202-2 gc202-4 gc202-5 vt8: milne land c. 10 95–90 gc1016-44 gc1016-45 gc1016-43 gc1016-42 c. 240 25°c/km unconformity ju ra ss ic si lu ri an c. 180 c8 c4 c1c2 vt9: kuhn ø c. 10 95–90 gc1077-22 gc1077-23 gc1077-24 gc1077-25 gc1077-26 c. 240 c8 c4 c1 25°c/km vt10: payer land c. 10 145–140 gc1077-27 gc1077-28 gc1077-29 gc1077-31 gc1077-32 c. 180 25°c/km gc1077-30 c. 5 95–90 c8 c3 c2c9 c4 vt11: stauning alper gc1077-18 gc1077-21 gc1077-20 25°c/km c. 10 c. 18095–90 c8 c2c4 vt12: liverpool land c. 10 gc1077-33 c. 18037–35 c8 c2c6 gc1077-34 gc1077-36 25°c/km gc1077-35 appendix 2.3 (continued) palaeotemperature constraints derived from afta data from vertical transects 1 to 17 plotted against elevation (a.s.l.). horizontal lines indicate range of palaeotemperatures estimated from afta data. arrows indicate lack of upper or lower limit for the estimated range. see appendix 2.2 for a summary of the geological constraints on the burial and exhumation history at the locations of these transects. locations in appendix 2.1. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 135 of 154 www.geusbul let in.org 0.5 1.0 0.0 1.5 timing (ma): episode: palaeotemperature (°c) 20 40 60 80 100 120 140 160 1800 2.0 el ev at io n (k m ) 0.5 1.0 0.0 1.5 timing (ma): episode: palaeotemperature (°c) 20 40 60 80 100 120 140 160 1800 2.0 el ev at io n (k m ) 0.5 1.0 0.0 1.5 timing (ma): episode: palaeotemperature (°c) 20 40 60 80 100 120 140 160 1800 2.0 el ev at io n (k m ) 0.5 1.0 0.0 1.5 timing (ma): episode: palaeotemperature (°c) 20 40 60 80 100 120 140 160 1800 2.0 el ev at io n (k m ) 0.5 1.0 0.0 1.5 timing (ma): episode: palaeotemperature (°c) ba se m en t ba se m en t ba se m en t pe rm o– tr ia ss ic l. c re t. pa la eo ge ne ju ra ss ic 20 40 60 80 100 120 140 160 1800 2.0 el ev at io n (k m ) ba se m en t pa le oc en e fig b2 continued2_jbl vt13: jameson land c. 10 gc1077-37 56–4837–35 c8 c5c6 gc1077-38 gc1077-39 25°c/km vt14: giesecke bjerge gc1077-46 gc1077-44 gc1077-43 25°c/km gc1077-45 c. 10 37–35 c8 c6 vt15: traill ø (e) gc1077-52 gc1077-55 gc1077-54 c. 10 37–35 c8 c6 c. 40°c/kmc. 40°c/km c. 5 c9 vt16: dombjerg c. 10 145–140 gc1077-91 gc1077-89 gc1077-92 c. 180 25°c/km gc1077-90 100–90c. 55 c8 c3 c2c4c5 vt17: clavering ø (ne) c. 10 c. 180 gc1077-96 gc1077-94 gc1077-97 c. 240 c8 c2 c1 25°c/km gc1077-93 gc1077-95 appendix 2.3 (continued) palaeotemperature constraints derived from afta data from vertical transects 1 to 17 plotted against elevation (a.s.l.). horizontal lines indicate range of palaeotemperatures estimated from afta data. arrows indicate lack of upper or lower limit for the estimated range. see appendix 2.2 for a summary of the geological constraints on the burial and exhumation history at the locations of these transects. locations in appendix 2.1. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 136 of 154 www.geusbul let in.org using the approach outlined in chapter 4 and methods documented in green et al. (2013), we have determined the range of palaeogeothermal gradients consistent with the palaeotemperature constraints from afta at the location of each vertical transect for palaeothermal episodes in which sufficient constraints are available (episodes c3, c4, c6 and c6 in appendix 2.4). extrapolation of allowed palaeogeothermal gradients makes it possible to determine the amounts of removed section required to explain the observed palaeotemperatures in each episode. note that determining amounts of removed section in this way is based on a number of assumptions, particularly that the palaeogeothermal gradient was linear throughout the entire succession when cooling from the palaeothermal peak began in each cooling episode. this is equivalent to assuming that the missing section and the preserved section comprised similar lithologies such that no significant contrasts in thermal conductivity existed, which may have resulted in a non-linear palaeotemperature profile (see fig. 32 of green et al. 2013). thermal conductivities measured in rocks of different types show wide variation from around 1 wm-1k-1 or less to over 5 wm-1k-1, related to lithology, water content, porosity, temperature and other factors (e.g. robertson 1988; clauser & huenges 1995; eppelbaum et al. 2014). if the cover rocks that were removed during an episode of exhumation were relatively unconsolidated sediments with lower thermal conductivities compared to the underlying basement rocks, then the palaeogeothermal gradient in the overlying section would have been higher. in that case, the amount of cover required to explain the miocene palaeotemperatures would be correspondingly reduced, perhaps by a factor of two, based on the range of typical thermal conductivities in the works cited above. the assumption of linearity will be invalid if the palaeotemperatures were caused by processes involving lateral or local introduction of heat, such as by confined fluid flow or igneous intrusion or where non-linear palaeotemperatures reflect palaeotopography. a palaeosurface temperature of 20°c was assumed in order to estimate amounts of removed section for the earliest cretaceous, the mid-cretaceous and the end-eocene episodes, decreasing to 10°c for the late miocene episode (based on zachos et al. 2001). changing the value of palaeosurface temperature is equivalent to a constant offset in the amount of missing section required to explain the observed palaeotemperatures. for example, increasing the palaeosurface temperature by 10°c, for a palaeogeothermal gradient of 25°c/km, would require a reduction of 400 m in the amount of removed section needed to explain the observed palaeotemperatures. for all episodes, amounts of additional (removed) section have been calculated with reference to the elevation of the highest local summit level. therefore, the reported values of removed section define how much was present above this level at the onset of each cooling episode. note, where palaeogene basalts are preserved at the summit (appendix 2.2), this provides a prime reference for calculating amounts of removed section for the end-eocene and younger events, as deposition required to explain these events must post-date deposition of the basalts. estimates of removed section for pre-palaeogene events are only possible at three locations (vt1 for the earliest cretaceous episode and vt9 and vt16 for the mid-cretaceous). in these cases, the estimate of removed section material relative to the top of the basalts represents a convenient approximation, as the base of the basalts represents the unconformity on which the additional section material was removed. of course, the base of the basalts would provide a more appropriate reference for calculation, but as the basalts are not too thick at the locations of the transects, the effect on the final calculation of removed section is only slight. early jurassic episode (c2): late jurassic palaeotemperature constraints from the vt3 transect suggest a high palaeogeothermal gradient (appendix 2.3). if the early jurassic episode was characterised by elevated heat flow, we would expect rapid increase in palaeotemperature with depth. therefore, this episode should be preferentially identified in samples close to sea level. this is not the case. instead, we see evidence for the older events still preserved at lower elevations close to present-day sea level. it seems more likely that this apparent anomaly results either from one of the palaeotemperatures being a statistical outlier or from offsets due to faulting. earliest cretaceous episode (c3): constraints on the earliest cretaceous episode are only available from vt1 (clavering ø). the allowed range of palaeogeothermal gradient (appendix 2.4) is quite well defined, but the range is broad, largely because constraints are available over an elevation range of only c. 800 m. similarly, the range of allowed values of additional section is very broad as this is magnified by the degree of extrapolation required by the analysis. the crossplot of removed section vs. palaeogeothermal gradient provides relatively well-defined ranges of removed section for specific values of palaeogradients. for example, a palaeogeothermal gradient of 25°c/km corresponds to a thickness of 1950 to 2200 m of additional section mechanisms of heating and cooling estimated from vertical transects palaeogeothermal gradients and amount of removed sections http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 137 of 154 www.geusbul let in.org vt1 clavering ø (w) summit elevation 1.1 km fig. b-3 vt16 dombjerg summit elevation 1.5 km vt17 clavering ø (ne) summit elevation 1.4 km vt3 hudson land (n) summit elevation 1.1 km vt5 geogr. society ø summit elevation 1.1 km vt8 milne land summit elevation 1.7 km vt6 traill ø (sw) summit elevation 1.4 km vt4 hudson land (e) summit elevation 1.5 km vt9 kuhn ø summit elevation 1.3 km vt13 jameson land summit elevation 1.2 km vt12 liverpool land summit elevation 1.2 km vt10 payer land summit elevation 1.3 km c3 c3 c8 c8 c8 c8 c6 c6 c6 c6 c6 c4 c4 early cretaceous episode late eocene episode maximum likelihood solution mid-cretaceous episode c4 0 10 20 30 geothermal gradient (°c/km) 40 50re m ov ed s ec ti on (k m ) 4 2 1 3 5 0 4 2 1 3 5 0 4 2 1 3 5 0 4 2 1 3 5 0 4 2 1 3 5 0 4 2 1 3 5 0 4 2 1 3 5 0 4 2 1 3 5 0 4 2 1 3 5 0 4 2 1 3 5 0 4 2 1 3 5 0 4 2 1 3 5 0 0 10 20 30 geothermal gradient (°c/km) 40 50 0 10 20 30 geothermal gradient (°c/km) 40 50 0 10 20 30 geothermal gradient (°c/km) 40 50re m ov ed s ec ti on (k m ) 0 10 20 30 geothermal gradient (°c/km) 40 50 0 10 20 30 geothermal gradient (°c/km) 40 50 0 10 20 30 geothermal gradient (°c/km) 40 50re m ov ed s ec ti on (k m ) 0 10 20 30 geothermal gradient (°c/km) 40 500 10 20 30 geothermal gradient (°c/km) 40 50re m ov ed s ec ti on (k m ) 0 10 20 30 geothermal gradient (°c/km) 40 50 0 10 20 30 geothermal gradient (°c/km) 40 50 0 10 20 30 geothermal gradient (°c/km) 40 50 c8 c8 c8 late miocene episode c8 c8 c8 c8 c8 c8 c6 appendix 2.4 comparison of allowed palaeogeothermal gradients vs. amount of removed section above and up to the present-day summit elevation as derived from afta data from 12 vertical transects (appendix 2.3). locations in appendix 2.1. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 138 of 154 www.geusbul let in.org that must have been present above the present-day summits of clavering ø to explain palaeotemperatures in episode c3. palaeogeothermal gradients in excess of 40°c/km are not allowed. so, for any reasonable value of palaeogeothermal gradient, the palaeotemperature constraints characterising this clearly require deposition and removal of around 2 km of post-carboniferous succession, prior to the onset of cooling in the interval 145 to 140 ma. on this basis, cooling is interpreted as due largely to exhumation. since palaeogene basalts and underlying sedimentary units are present at the summit of this transect, erosional removal of all the additional section must have been accomplished by around the end of the cretaceous. mid-cretaceous episode (c4): constraints on the mid-cretaceous episode are only available from transects vt9 (kuhn ø) and vt16 (dombjerg). appendix 2.4 shows that the ranges of allowed values of palaeogeothermal gradients for each transect are well defined, despite the relatively narrow intervals over which constraints are available in these transects. however, the range of allowed removed sections is magnified by the degree of extrapolation required to determine the amount of missing section. nevertheless, the crossplot of removed section vs. palaeogeothermal gradient provides relatively well-defined ranges of removed section for specific values of palaeogradient. these plots show that a palaeogeothermal gradient of 25°c/km corresponds to a thickness of 1550 to 1850 m of additional section for vt9 and 1800 to 2600 m for vt16. at both locations, palaeogeothermal gradients in excess of 35°c/km are not allowed, so for any reasonable value of palaeogeothermal gradient, the palaeotemperature constraints characterising this episode, clearly require deposition and removal of around 2 km of section, prior to the onset of cooling in the interval 95 to 90 ma. since palaeogene basalts and underlying sedimentary units are present at the summit of this transect, erosional removal of all the additional section must have been accomplished by around the end of the cretaceous. end-eocene episode (c6): constraints on the end-eocene episode are available from transects vt3, vt4, vt6 vt12 and vt13 (appendix 2.4). results from vt4 (hudson land) provide the best definition of the palaeotemperature profile for this episode, in terms of numbers of samples and the vertical extent of samples. the plot of removed section vs. palaeogeothermal gradient for vt4 shows a relatively narrow range of allowed values between c. 20 and 40°c/km. results from vt6 are also quite well defined, while the ranges of allowed palaeogradients for other transects (e.g. vt3, vt12 and vt13) are wider. nevertheless, the crossplots of removed section vs. palaeogeothermal gradient provide relatively well-defined ranges of removed section for specific values of palaeogradient for all five transects. for vt4, the best-fit palaeogeothermal gradient of 30°c/km corresponds to around 1.65 km of additional section required to explain the end-eocene palaeotemperatures, and for any value of palaeogradient within the allowed range, km-scale additional section is required. since no palaeogene basalts are present at this location, some of this additional section could represent basalt. results from transect vt6 tend to favour higher values of palaeogeothermal gradient while others generally tend to favour values around 20 to 30°c/km. while higher palaeogradients require lower amounts of additional section, for any reasonable value of palaeogradient, km-scale burial and exhumation are required to explain the end-eocene palaeotemperature constraints derived from afta data in all five transects. late miocene episode (c8): of all the palaeothermal episodes, the late miocene is most clearly expressed over the majority of vertical transects (appendix 2.4), and it is the most widely recognised across the entire region (map c8 in fig. 14). this episode appears to dominate the development of the present-day continental margin. constraints on the late miocene episode are available from all transects except vt7, vt14 and vt15. transect vt7 is excluded because of uncertainty surrounding exact sample elevations, vt14 because samples are only available over a narrow depth interval and vt15 because the palaeothermal constraints are too broad. the range of allowed values of additional section is very broad for all transects, with no upper limit being defined in many cases. but the crossplots of removed section vs. palaeogeothermal gradient in most cases provide relatively well-defined ranges of removed section for specific values of palaeogradient. the more well-constrained results from these transects tend to favour palaeogeothermal gradients in the range 20 to 30°c/km, while results from many transects are not consistent with palaeogradients above c. 40°c/km. palaeogradients around 20 to 30°c/km generally correspond to amounts of removed section between 1 and 2 km. since these calculations have been carried out with respect to local summit elevations, the results shown here require removal of a considerable amount of additional section over the last 10 myr. and since early eocene basalts are present at many of the summits, the additional section required to explain the late miocene palaeotemperatures, must represent additional eocene basalts and younger sediments deposited in the time interval between eruption of the basalts and the onset of exhumation at c. 10 ma. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 139 of 154 www.geusbul let in.org comparison of transects for the mesozoic and end-eocene events appendix 2.5 shows palaeotemperatures characterising the middle triassic (c1), early jurassic (c2), earliest cretaceous (c3) and end-eocene (c6) palaeothermal episodes in each of the vertical transects plotted against a common axis of sample elevation for each episode. results from each transect and for each episode can generally be described by profiles with gradients of c. 25°c/km, although only a small number of transects provide any real constraints on the palaeogeothermal gradient. the figure also illustrates how the earlier episodes are only identified at locations where the respective maximum palaeotemperatures are sufficiently higher than those in later episodes. again, we emphasise that the apparent high gradient defined by the early jurassic palaeogeothermal gradient from vt3 stands out as anomalous when compared with other results. palaeotemperatures during the earliest cretaceous (c3) episode highlight some major differences between transects. in particular, palaeotemperatures in samples from vt4 and vt16 are offset to higher elevations by between 500 and 1000 m from those in samples from vt1 and vt3, while the single value from vt10 plots at an intermediate position. such offsets clearly demand major differential exhumation between these locations during or after the earliest cretaceous cooling episode. appendix 2.5 provides a comparison of palaeotemperatures in the end-eocene (c6) episode at eight vertical transects where this episode is identified by afta. some clear differences are evident. vt3 palaeotemperatures are lower than those from vt4, vt11, vt11, vt12 and vt13, while those from vt6, vt14 and vt15 are higher than the other five. palaeotemperature constraints from most of the transects can be described by palaeogradients around 25°c/km. vt6 defines a distinctly higher palaeogeothermal gradient, in line with vt14 and vt15. these higher palaeogradients can be explained in terms of their proximity to the middle eocene – early oligocene intrusions at myggbukta (vt14) and traill ø (vt6 and vt15). http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 140 of 154 www.geusbul let in.org c3 145–140 ma c2 c. 180 mac1 c. 240 ma c6 37–35 ma fig b-4_jbl adjacent to intrusions el ev at io n (k m ) 0.5 1.0 0.0 1.5 palaeotemperature (°c) 20 40 60 80 100 120 140 160 1800 2.0 el ev at io n (k m ) 0.5 1.0 0.0 1.5 palaeotemperature (°c) 20 40 60 80 100 120 140 160 1800 2.0 el ev at io n (k m ) 0.5 1.0 0.0 1.5 palaeotemperature (°c) 20 40 60 80 100 120 140 160 1800 2.0 el ev at io n (k m ) 0.5 1.0 0.0 1.5 palaeotemperature (°c) 20 40 60 80 100 120 140 160 1800 2.0 vt1: clavering ø vt1: clavering ø vt3: hudson land vt3: hudson land vt3: hudson landvt4,4a: hudson land vt4,4a: hudson land vt4,4a: hudson landvt10: payer land vt10: payer land vt11: stauning alper vt12: liverpool land vt12: liverpool land vt16: dombjerg vt16: dombjerg vt8: milne land vt9: kuhn ø vt17: ne clavering ø vt13: jameson land vt14: giesecke bjerge vt15: eastern traill ø vt6: traill ø appendix 2.5 comparison of palaeotemperature constraints for four episodes derived from afta plotted against elevation (a.s.l.), in samples from vertical transects: c1 middle triassic, c2 early jurassic, c3 earliest cretaceous and c6 end-eocene. constraints for each transect are colour-coded as indicated. results for the end-eocene episode fall into three dominant groupings for low, medium and high palaeotemperatures for a particular elevation. highest values are all close to middle eocene – early oligocene intrusions, and these results suggest high palaeogeothermal gradients (45°c/km for vt6, vt14 and vt15) in contrast to lower profiles (25°c/ km). locations in appendix 2.1. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 141 of 154 www.geusbul let in.org appendix 3 thermal-history diagrams appendices 3.1 to 3.13 illustrate the thermal history at summit level above a vt or a group of samples, based on palaeothermal constraints from afta (appendix 2), the geological record (fig. 3; appendix 2.2) and stratigraphic landscape analysis (chapter 3). the diagrams are constructed in the simplest possible manner to provide an overview of the available constraints. note that the time axis has two different scales before and after 65 ma. we have assumed that 8 of the 10 cooling episodes affected most of the study area even if the afta data only resolve a minor number of episodes at a given location (table 1). location of vts and plots of palaeothermal constraints vs. elevation for each transect are shown in appendices 2.1, 2.3. amounts of removed section vs. palaeogeothermal gradient are shown for each transect in appendix 2.4. a map of the extent of the ups and lps is shown in fig. 9. palaeotemperature constraints are projected to the summit level for a palaeogeothermal gradient of 25°c/ km for each cooling episode recognised in afta data for the samples along each transect unless otherwise indicated. the ranges of palaeotemperatures for the summit level in different episodes are plotted as boxes with a height corresponding to the allowed range; typically 10–20°c. to enhance the readability of the diagram, the width of the boxes is fixed in centimetres and does thus not reflect the constraints on the onset of cooling from the samples along the transect. boxes are extended below the x-axis of the diagram where the palaeothermal constraints place only a lower limit on the palaeotemperature (e.g. >110°c). coloured vertical bars define the onset of regional cooling episodes (table 1). red triangles mark constraints from geology; e.g. the time of sediment deposition or extrusion of basalts for rocks that outcrop at the summits. at that time, the rock (often basement) below these outcrops was exposed at the surface and thus had a palaeotemperature of 10–20°c depending on the climate. the horizontal blue arrows mark the time available for the formation of the ups between the onset of uplift and cooling in the end-eocene (c6) and the inferred onset of miocene reburial which is arbitrarily taken to start at 20 ma. if instead, the lps coincides with the summit level, and we assume that the lps was incised 500 m below a previously existing ups, then the rocks at summit level were cooled to about 20°c (c. 10°c + 0.5 km per 25°c/km). the horizontal red arrows mark the time available for the formation of the lps between the onset of uplift and cooling in the late miocene (c8) and the onset of uplift in the early pliocene (c9). for simplicity, the cooling path between the late miocene palaeotemperature and the present-day surface temperature (arbitrarily shown as 0°c) is shown as a straight line. interpretations are provided in the appendix captions. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 142 of 154 www.geusbul let in.org upper carboniferous sandstones heating below upper carboniferous to middle triassic sediments ?? ? basalts ? ? vt10 hotter than vt1 ? heating below eocene to mid-miocene rocks ? appendix 3.1 vt1. summits of western clavering ø. local summits 1.2 km (close to the lps). upper carboniferous sandstones are present in the summits on western clavering ø, just below palaeogene basalts. the sandstones were therefore exposed at the surface in the carboniferous and in the early eocene. this implies that late miocene c8 palaeotemperatures are due to burial below palaeogene basalts and post-basalt sediments deposited between c. 55 and 10 ma. middle triassic c1 palaeotemperatures are due to burial below upper carboniferous to middle triassic sediments. c2 palaeotemperature s are not resolved for vt1, but must have been between the c1 and c3 values. we suggest that c2 palaeotemperatures for vt1 must be lower than those resolved for vt10, payer land (i.e. <100°c). http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 143 of 154 www.geusbul let in.org ? local summits 1.2 km (near lps) middle jurassic sediments above triassic sample basalt ? heating below middle jurassic to lowermost cretaceous sediments heating below eocene to mid-miocene rocks ? appendix 3.2 vt2. summits at north-west hold with hope. local summits 1.2 km (close to the lps). middle jurassic sandstone units present near the summits of vt2 were deposited on triassic sediments at the surface at a temperature of c. 20°c. afta data show that a sample of triassic sandstone was heated to 60–80°c during the earliest cretaceous c3 episode. this implies the deposition of a kilometre-thick cover of middle jurassic to lowermost cretaceous sediments prior to the onset of episode c3. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 144 of 154 www.geusbul let in.org heating to >100°c below upper carboniferous to middle triassic sediments ?? ? ? palaeogene basalts in the summits of clavering ø ? upper carboniferous sandstones ? appendix 3.3 vt3. summits of hudson land (n). local summits 1.3 km (close to the lps). ups at 1.8 km. all vt3 samples are of upper carboniferous sandstones and each sample was therefore at surface temperature during deposition. no palaeogene basalts are present in the summits above vt3, but basalts are present in the summits of clavering ø and on hold with hope. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 145 of 154 www.geusbul let in.org ?? ? ? palaeogene basalts in the summits of clavering ø ? upper carboniferous sandstones in the summits of vt3 ? ? appendix 3.4 vt4. summits of hudson land (s). local summits 1.3 km (close to the lps). ups at 1.8 km. all vt4 samples are of basement rocks, but upper carboniferous sandstones are present nearby (vt3). heating to temperatures above 100°c, characteristic of episode c1, reflects burial below carboniferous to middle triassic sediments. no palaeogene basalts are present in the summits above vt4, but basalts are present in the summits of clavering ø and hold with hope. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 146 of 154 www.geusbul let in.org ? ? ? basalts middle jurassic sediments above triassic sample heating below middle jurassic to lower cretaceous sediments ? appendix 3.5 vt5. summits on geographical society ø. local summits 1.1 km (close to the lps). triassic sandstone units were deposited at the surface at a temperature of c. 20°c, and now crop out below middle jurassic sediments. afta data show that the triassic sample was heated to almost 90°c during the mid-cretaceous c4 episode. this implies the deposition of a kilometre-thick cover of middle jurassic to lower cretaceous sediments prior to the onset of episode c4. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 147 of 154 www.geusbul let in.org ? triassic–cretaceous sediments ? ? appendix 3.6 vt5. summits on geographical society ø. local summits 1.1 km (close to the lps). triassic sandstone units were deposited at the surface at a temperature of c. 20°c, and now crop out below middle jurassic sediments. afta data show that the triassic sample was heated to almost 90°c during the mid-cretaceous c4 episode. this implies the deposition of a kilometre-thick cover of middle jurassic to lower cretaceous sediments prior to the onset of episode c4. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 148 of 154 www.geusbul let in.org middle jurassic sandstones ? ? ? basalts heating below middle jurassic to mid-cretaceous sediments heating below eocene to mid-miocene rocks ? appendix 3.7 vt9. summits of kuhn ø. local summits 1.2 km (close to the lps). kuhn ø is a tilted late jurassic – early cretaceous fault block where middle jurassic sediments onlap basement on the western side of the island. the basement rocks in the summits were thus near the surface in the middle jurassic, and the mid-cretaceous c4 palaeotemperatures therefore correspond to burial below a thick cover of middle jurassic – mid-cretaceous sediments. palaeogene basalts are present in the summits. thus, the basement rock was near the surface in the early eocene. this implies that c8 palaeotemperatures at 10 ma are due to burial below basalts and post-basalt sediments deposited between c. 55 and 10 ma. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 149 of 154 www.geusbul let in.org ??? ? ? ? ? upper carboniferous sandstones in the summits of vt1 appendix 3.8 vt10. summits of payer land. local summits 1.3 km (close to the lps). on payer land there are no deposits present that constrain the phanerozoic development. however, results from the nearby vt1 on eastern clavering ø show that much of the heating prior to c2 cooling must have involved burial below upper carboniferous sediments. one sample at low elevation provides constraints on the pliocene palaeotemperature. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 150 of 154 www.geusbul let in.org ?? burial of the sub-permian peneplain ? ? ? sub-permian peneplain west of liverpool land ? ? ? appendix 3.9 vt12. summits of liverpool land. local summits 1.2 km (close to the lps). a westward tilted, sub-permian peneplain extends across the western flanks of liverpool land and below upper permian and triassic sediments in the jameson land basin. the basement rocks in the summits were thus not far from the surface in the mid-permian. palaeotemperatures in the early jurassic c2 episode are in excess of 100°c, implying that the summits at that time were buried below a thick cover of upper permian – lower jurassic sediments. heating to high palaeotemperatures in the late eocene was probably partly due to eocene burial of the margin. likewise, heating in the late miocene also reflects miocene reburial of the margin. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 151 of 154 www.geusbul let in.org ??? ? basalt heating below eocene to mid-miocene rocks ? appendix 3.10 vt16. summits of dombjerg. local summits 1.4 km (close to the lps). palaeogene sandstones and basalts are present in the summits of dombjerg. the basement rocks just below the palaeogene deposits were thus exposed at the surface in the earliest eocene. this implies that the c8 palaeotemperatures at 10 ma were due to burial below post-basalt sediments deposited between c. 55 and 10 ma. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 152 of 154 www.geusbul let in.org upper permian limestones ?? ? basalts ? vt16 hotter than vt17 heating below eocene to mid-miocene rocks ? appendix 3.11 vt17. summits of north-east clavering ø. local summits 1.2 km (close to the lps). upper permian carbonates crop out in the summits of clavering ø where basement samples from adjacent locations were heated to a middle triassic c1 palaeotemperature in excess of 110°c. heating at locations where carbonates are not present, was caused by burial below a thick cover of upper permian to middle triassic sediments plus an unknown column of basement rocks. palaeogene sandstones and basalts are present in the summits of vt17. basement rocks just below the palaeogene deposits were thus exposed at the surface in the early eocene. this implies that late miocene c8 palaeotemperatures were due to burial below basalts and post-basalt sediments deposited between c. 55 and 10 ma. c3 palaeotemperature in vt17 is not resolved, but it must have been lower than c2 in this transect (i.e. <80°c). this would be substantially lower than that resolved for c3 in vt16 (dombjerg, appendix 3.10). these differences between vt16 and vt17 were levelled out prior to the eruption of the palaeogene basalts (present in the summits at both locations at about the same elevation); possibly during the mid-cretaceous c4 episode. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 153 of 154 www.geusbul let in.org ?? ? ? ? ? ? middle – upper jurassic sediments lower cretaceous sediments heating below lower cretaceous to mid-cretaceous sediments ? appendix 3.12 store koldewey at sea level. basement sample gc1077-16 cooled below 110°c in the late carboniferous c0 episode. this sample and a sample of middle jurassic sandstone cooled from about 80°c in the mid-cretaceous c4 episode. both samples are adjacent to lower cretaceous sediments at outcrop, and this implies deposition of a kilometre-thick lower cretaceous – mid-cretaceous cover prior to onset of uplift and erosion in episode c4. the samples were heated to c. 60°c prior to cooling in the late miocene c8 episode. http://www.geusbulletin.org japsen et al. 2021: geus bulletin 45(2). 5299. https://doi.org/10.34194/geusb.v45.5299 154 of 154 www.geusbul let in.org ? ? ? ? ? ? upper carboniferous sediments middle jurassic sediments heating below middle jurassic to mid-cretaceous sediments ? appendix 3.13 germania land at sea level. basement samples gc1077-10 and -11 cooled below 105°c in the late carboniferous c0 episode. assuming this episode corresponds to the unconformity between basement and an upper carboniferous outlier on germania land (piasecki et al. 1994), then the cooling from these palaeotemperatures to surface conditions represents kilometre-scale exhumation prior to late carboniferous deposition. heating of samples gc1077-9, -10 and -11 to palaeotemperatures of c. 100°c or more in the middle triassic c1 episode reflects burial below a thick upper carboniferous – middle triassic cover. middle jurassic sedimentary outliers also rest directly on basement on germania land (bojesen-koefoed et al. 2012), again implying kilometre-scale exhumation between the onset of middle triassic cooling and middle jurassic deposition. samples gc1017-9 and -11 reached a palaeotemperature of c. 75°c prior to mid-cretaceous cooling and thus reflect significant burial during middle jurassic to mid-cretaceous. http://www.geusbulletin.org episodic burial and exhumation in north-east greenland before and after opening of the north-east at abstract 1. introduction 2. post-caledonian geology 2.1 carboniferous-cretaceous cover and unconformities 2.2 cenozoic cover and unconformities 2.3 palaeogene intrusives 2.4 sediments and unconformities offshore north-east greenland 3. stratigraphic landscape analysis 3.1 preand post-basalt peneplains 3.2 relative denudation chronology 4. apatite fission-track analysis data 4.1 afta and vr principles 4.2 previous apatite fission-track studies 4.3 new afta and vr data 4.4 thermal-history solutions from afta 4.5 regional variation in the magnitude of palaeothermal episodes 4.5.1 late carboniferous episode (c0) 4.5.2 middle triassic episode (c1) 4.5.3 early jurassic episode (c2) 4.5.4 earliest cretaceous episode (c3) 4.5.5 mid-cretaceous episode (c4) 4.5.6 early eocene episode (c5) 4.5.7 end-eocene episode (c6) 4.5.8 early miocene episode (c7) 4.5.9 late miocene episode (c8) 4.5.10 early pliocene episode (c9) 4.6 palaeogeothermal gradients inferred from afta in vertical transects 4.6.1 principles of the method 4.6.2 palaeotemperature profiles in vertical transects 4.6.3 estimates of palaeogeothermal gradients 4.6.4 summary 5. integration of afta results with geological evidence 5.1 late carboniferous episode (c0) 5.2 middle triassic episode (c1) 5.3 early jurassic episode (c2) 5.4 earliest cretaceous episode (c3) 5.5 mid-cretaceous episode (c4) 5.6 early eocene episode (c5) 5.7 end-eocene episode (c6) 5.8 early miocene episode (c7) 5.9 late miocene episode (c8) 5.10 the early pliocene episode (c9) 5.11 comparison with results from adjacent areas of east greenland 6. development of the continental margin of north-east greenland after break-up 6.1 eocene burial and exhumation 6.1.1 evidence for eocene subsidence/burial and uplift/exhumation 6.1.2 extent of eocene cover and of end-eocene exhumation 6.1.3 end-eocene differential, vertical movements 6.2 miocene burial and exhumation 6.2.1 development of the present-day topography within the last 10 myr 6.2.2 late miocene differential, vertical movements 6.3 cenozoic events in the traill ø area 6.4 stauning alper 6.5 a stepped topography formed after three phases of uplift and incision 7. palaeozoic-mesozoic thermal histories 7.1 thermal-history reconstructions, store koldewey 7.1.1 samples gc1017-16, -102, crystalline basement, near sea level 7.1.2 sample gc1017-83, crystalline basement, near sea level 7.2 thermal-history reconstructions of pre-mesozoic terrains 7.2.1 sample gc1016-62, crystalline basement, hudson land (123 m a.s.l.) 7.2.2 sample gc1016-55, carboniferous sediments, clavering ø (52 m a.s.l.) 7.2.3 sample gc1016-9, devonian sediments, ella ø (1 m a.s.l.) 7.2.4 sample gc1016-45, ordovician intrusion, milne land (458 m a.s.l.) 7.3 thermal-history reconstructions of mesozoic terrains 7.3.1 sample gc1016-69, triassic sediments, hold with hope (94 m a.s.l.) 7.3.2 sample gc522-7, jurassic sediments, wollaston forland (150 m a.s.l.) 7.3.3 sample gc522-10, jurassic sediments, traill ø (230 m a.s.l.) 7.3.4 sample gc202-11, triassic sediments, south-eastern jameson land (150 m a.s.l.) 7.4 palaeozoic-mesozoic removed covers 7.4.1 cover prior to exhumation in the middle triassic episode (c1) 7.4.2 cover prior to exhumation in the early jurassic episode (c2) 7.4.3 cover prior to exhumation in the earliest cretaceous episode (c3) 7.4.4 cover prior to exhumation in the mid-cretaceous episode (c4) 7.5 fault movements in the clavering ø area 8. hydrocarbon prospectivity 8.1 store koldewey 8.2 exhumed oil accumulations on traill ø 8.3 maturation studies of the jameson land basin 9. episodes of uplift and erosion in greenland and scandinavia 9.1 phases of exhumation prior to break-up of pangaea 9.1.1 late carboniferous episode (c0) 9.1.2 middle triassic episode (c1) 9.1.3 early jurassic episode (c2) 9.2 phases of exhumation after break-up of pangaea and before the opening of the north-east atlantic 9.2.1 earliest cretaceous episode (c3) 9.2.2 mid-cretaceous episode (c4) 9.3 phases of exhumation after opening of the north-east atlantic 9.3.1 end-eocene episode (c6 9.3.2 early miocene uplift and erosion in southern scandinavia 9.3.3 late miocene episode (c8) 9.3.4 early pliocene episode (c9) 9.4 development of the modern topography in greenland and scandinavia 10. discussion 10.1 imposing geological constraints on thermal-history reconstructions 10.2 tectonic uplift vs erosional unloading 10.3 development of the greenland ice sheet 10.3.1 late neogene onset of the formation the greenland ice sheet 10.3.2 continental ice in east greenland at the eocene-oligocene transition 10.4 episodic tectonic processes along margins and their hinterland 11. conclusions 11.1 palaeothermal episodes 11.2 development of present-day topography 11.3 fault offsets 11.4 onshore-offshore correlation 11.5 regionally synchronous episodes 11.6 non-synchronous events 11.7 episodic burial and exhumation acknowledgements additional information references appendix 1 thermal history constraints from afta data in individual samples and definition of region thermal history interpretation of afta data: principles appendix 1.1 illustration of thermal history solutions extracted from afta. a full description is pr results from north-east greenland appendix 1.2 sample details. a full description is provided in the table caption. appendix 1.3 afta data. a full description is provided in the table caption. appendix 1.4 thermal history solutions. a full description is provided in the table caption. appendix 1.5 maximum palaeotemperatures. a full description is provided in the table caption. timing constraints derived from afta data in individual samples appendix 1.6.1 region 1 timing constraints. a full description is provided in the figure caption. appendix 1.6.2 region 2 timing constraints. a full description is provided in the figure caption. appendix 1.6.3 region 3 timing constraints. a full description is provided in the figure caption. appendix 1.6.4 region 4 timing constraints. a full description is provided in the figure caption. appendix 1.6.5 region 5 timing constraints. a full description is provided in the figure caption. appendix 1.6.6 region 6 timing constraints. a full description is provided in the figure caption. appendix 1.7 thermal history interpretations. a full description is provided in the table caption. appendix 2 results from vertical transects individual vertical transects appendix 2.1 location of vertical transects. a full description is provided in the figure caption. appendix 2.2 geological constraints on the burial and exhumation history. a full description is prov appendix 2.3 palaeotemperature constraints from afta. a full description is provided in the figure c mechanisms of heating and cooling estimated from vertical transects palaeogeothermal gradients and amount of removed sections appendix 2.4 palaeogeothermal gradients vs. amount of removed section. a full description is provide comparison of transects for the mesozoic and end-eocene events appendix 2.5 palaeotemperature constraints for four episodes. a full description is provided in the appendix 3 thermal-history diagrams appendix 3.1 thermal history diagram for vt1. a full description is provided in the figure caption. appendix 3.2 thermal history diagram for vt2. a full description is provided in the figure caption. appendix 3.3 thermal history diagram for vt3. a full description is provided in the figure caption. appendix 3.4 thermal history diagram for vt4. a full description is provided in the figure caption. appendix 3.5 thermal history diagram for vt5. a full description is provided in the figure caption. appendix 3.6 thermal history diagram for vt6. a full description is provided in the figure caption. appendix 3.7 thermal history diagram for vt9. a full description is provided in the figure caption. appendix 3.8 thermal history diagram for vt10. a full description is provided in the figure caption. appendix 3.9 thermal history diagram for vt12. a full description is provided in the figure caption. appendix 3.10 thermal history diagram for vt16. a full description is provided in the figure caption appendix 3.11 thermal history diagram for vt17. a full description is provided in the figure caption appendix 3.12 thermal history diagram for store koldewey at sea level. a full description is provide appendix 3.13 thermal history diagram for germania land at sea level. a full description is provided figures fig. 1: location of study area in north-east greenland. a full description is provided in the figure fig. 2: elevation and place names of the study area in north-east greenland. a full description is p fig. 3: geological map of north-east greenland. a full description is provided in the figure caption fig. 4: lithostratigraphic scheme for north-east greenland. a full description is provided in the fi fig. 4: lithostratigraphic scheme for north-east greenland. a full description is provided in the fi fig. 5 radiometric ages of cenozoic igneous rocks along the east greenland margin a full description fig. 6 seismic profiles north-east greenland. a full description is provided in the figure caption. fig. 7 photo of milne land (background of the photo) and profile. a full description is provided in fig. 8 photo of flyverfjord and nordvestfjord and a 3d elevation model. a full description is provid fig. 9 map showing extent and elevation of the upper and lower planation surfaces. a full descriptio fig. 10 schematic illustration of the afta method. a full description is provided in the figure capt fig. 11 afta and vr sample locations. a full description is provided in the figure caption. fig. 11 afta and vr sample locations. a full description is provided in the figure caption. fig. 12 fission-track ages in individual outcrop samples. a full description is provided in the figu fig. 13 scatter graph showing relationship between mean track length and fission-track age. a full d fig. 14 maps of palaeotemperatures derived from afta data. a full description is provided in the fig fig. 14 maps of palaeotemperatures derived from afta data. a full description is provided in the fig fig. 14 maps of palaeotemperatures derived from afta data. a full description is provided in the fig fig. 15 map of thermal-history solutions derived from afta and vr data for store koldewey. a full de fig. 16 maps of palaeotemperatures from afta and vr around clavering ø for episodes c3, c4, c6 and c fig. 17 scatter graph comparing palaeogeothermal gradients in four palaeothermal episodes (c3, c4, c fig. 18 regional cooling episodes overlain on a lithostratigraphic scheme for north-east greenland. fig. 18 regional cooling episodes overlain on a lithostratigraphic scheme for north-east greenland. fig. 19 jurassic stratigraphy in the wollaston forland area along a n–s profile and timing of the pr fig. 20 summary of the main tectonic events and trends in jurassic – early cretaceous basin evolutio fig. 21 north–south extent of cenozoic cooling events identified from afta data along the east green fig. 22 late miocene c8 palaeotemperature constraints from afta. a full description is provided in t fig. 23 timing of post-devonian cooling episodes in east greenland. a full description is provided i fig. 24 geological map of the northern part of the study area. a full description is provided in the fig. 25 schematic showing projected surface elevation corresponding to palaeotemperatures for a pala fig. 26 map showing common projected surface elevation for end-eocene c6 episode. a full description fig. 27 geological map of blosseville kyst and end-eocene c6 palaeotemperatures in three samples. a fig. 28 thermal-history schematic for the summits of hudson land. a full description is provided in fig. 29 photo showing the contact between basement and palaeogene basalts along the storelv fault, h fig. 30 map showing common projected surface elevations for the late miocene episode. a full descrip fig. 31 on the left, a scatter graph showing constraints on late miocene c8 palaeotemperatures from fig. 32 a geological map and graphs of palaeotemperature constraints from afta in samples from four fig. 33 six maps showing cenozoic uplift events and relief formed by denudation in the study area. a fig. 34 block diagrams illustrating the post-rift development of the east greenland margin since cen fig. 35 geological map and thermal history diagrams in caledonian basement samples from store koldew fig. 36 geological map and thermal history diagrams for samples of palaeozoic age. a full descriptio fig. 37. geological map and thermal history diagrams from samples of mesozoic age. a full descriptio fig. 38 geological map and outline of middle jurassic to upper cretaceous deposits. a full descripti fig. 39 geological map and numerous graphs of palaeotemperature constraints from afta. a full descri fig. 40 elevation profile from milne land to liverpool land. a full description is provided in the f fig. 41 timing of regional, post-devonian episodes of uplift and erosion. a full description is prov fig. 42 conflicting interpretations of the post-caledonian development in north-east greenland. a fu table 1 intervals defining the onset of episodes of cooling. a full description is provided in the t data article | short schullehner 2022: geus bulletin 49. 8319. https://doi.org/10.34194/geusb.v49.8319 1 of 6 danish water supply areas and their links to water production facilities: an open-access data set jörg schullehner*1,2,3  1geological survey of denmark and greenland, aarhus, denmark. 2department of public health, environment, occupation and health, aarhus university, aarhus, denmark. 3danish big data centre for environment and health (bertha), aarhus university, aarhus, denmark. abstract this data set establishes the missing link between drinking-water quality monitoring data at the water production facility level in the danish national geodatabase jupiter and supply areas. water supply areas (wsas) were collected at municipality level, digitised and linked to the waterworks they are supplied by. infrastructural changes between 1978 and 2019 were taken into account by allowing wsa polygons to change over time. the number of active wsas decreased from 3172 in 1978 to 2602 in 2019. the data set consists of longitudinal wsa polygons and a table linking wsas to the water production facility identification in the jupiter database, allowing the estimation of current and historical drinking-water quality across denmark. in combination with the danish address register and the civil registration system, this data set allows exposure assessments of drinking-water quality at high spatiotemporal resolution for the entire danish population. therefore, this data set is an essential part of studying health effects of drinking-water quality in epidemiological research in denmark. *correspondence: jorg.schullehner@ph.au.dk received: 07 apr 2022 accepted: 17 may 2022 published: 28 june 2022 keywords: drinking water, monitoring, exposure assessment, epidemiology, denmark abbreviations: gis: geographic information system pfas: perand polyfluoroalkyl substances pin: personal identification number plantid: water production facility identifier wsa: water supply area wsaid: wsa identifier geus bulletin is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: adam hambly (technical university of denmark) reviewed by: debbie white (british geological survey, uk), martin rygaard (technical university of denmark). funding: see page 6 competing interests: see page 6 additional files: see page 6 tabular abstract geographical coverage denmark temporal coverage 1978–2019 subject(s) engineering and environmental geology data format(s) analysed spatial data (shapefile) and linkage table (text file) sample collection and analysis water supply area polygons collected and digitised from various sources, mainly at municipality level. waterworks and drinking-water quality data from jupiter parameters water supply area polygons with start and end year of activity, linkage of water supply area polygons (wsaid) to plantid (anlaegid) in jupiter. related publications schullehner & hansen 2014; voutchkova et al. 2015, 2021; schullehner et al. 2017, 2018, 2019, 2020; wodschow et al. 2018, 2021; coffman et al. 2021; stayner et al. 2021, 2022; thomsen et al. 2021; thygesen et al. 2021; ebdrup et al. 2022; richter et al. 2022; skaarup et al. 2022; theisen et al. 2022. potential application(s) for these data linkage of water quality registered at production facility level in jupiter to supply zones, households. allows estimation of drinkingwater quality exposure of the entire danish population through geocoded residential history and thus linkage to epidemiological studies on health effects of drinking-water quality. introduction data on drinking-water quality are relevant for a variety of stakeholders, including authorities, interest groups, plumbers, the general public and researchers. in denmark, drinking-water quality is monitored systematically, and results are registered in a nationwide, publicly accessible geodatabase called jupiter (hansen & pjetursson 2011; geus 2022). jupiter holds drinking water-related information including administrative information on https://doi.org/10.34194/geusb.v49.8319 https://orcid.org/0000-0002-1153-6885 mailto:jorg.schullehner@ph.au.dk schullehner 2022: geus bulletin 49. 8319. https://doi.org/10.34194/geusb.v49.8319 2 of 6 www.geusbul let in.org drinking-water production facilities, technical construction and geological description of boreholes, water abstraction volumes, and drinking-water quality samples and their analytical results. additionally, jupiter holds information on the location of the water production facilities and their boreholes. however, jupiter does not hold any information on the areas that water production facilities supply. this missing link is crucial for a range of tasks, including monitoring of drinking water at population level and estimation of historical drinking-water exposure for use in epidemiological studies on health effects of drinkingwater quality. to overcome these limitations, efforts have been made to establish this missing link by collecting and digitising the danish water supply areas (wsas) and linking them to water production facilities in jupiter (schullehner & hansen 2014; schullehner et al. 2017). this data set has already been used in many epidemiological studies (schullehner et al. 2018, 2019, 2020; coffman et al. 2021; stayner et al. 2021, 2022; thomsen et al. 2021; thygesen et al. 2021; wodschow et al. 2021; ebdrup et al. 2022; richter et al. 2022; theisen et al. 2022) and nationwide drinking-water quality mapping and modelling (wodschow et al. 2018; voutchkova et al. 2021, skaarup et al. 2022) and has been continuously developed to include the historical wsas from 1978 to 2019. here, this data set is described and published as an open-access data set. data collection setting danish drinking-water supply has several distinctive characteristics: it is based on groundwater, often with limited treatment (typically aeration and sand filtration), without chlorination and decentralised: around 2600 public waterworks supply a population of around 5.8 million (schullehner & hansen 2014; ministry of environment of denmark 2022). groundwater for drinking production is mainly sourced from aquifers in limestone/ chalk and quaternary and miocene sandy deposits. drinking-water quality monitoring is based on the implementation of the eu drinking water directive (european council 1998) in the ministerial order on water quality and control of water supply facilities (ministry of environment of denmark 2021). historically, the ministerial order defined public supplies (in danish:  almene vandværker) as serving ten or more households (schullehner & hansen 2014). this definition is applied in the presented data set that focuses on public supplies. the ministerial order defines monitoring programmes, including lists of quality parameters to be monitored and the sampling frequency depending on water production volumes. the monitoring programme has changed over time, resulting in data breaks, which need to be considered. according to the ministerial order, certified laboratories must take and analyse drinking-water samples, and results of both raw and finished treated drinking-water analyses and annual water production volumes are to be reported to the publicly accessible jupiter database. jupiter currently holds information on more than 370  000 boreholes, more than 40  000 of which are linked to water production facilities as well as around 60  000 water production facilities; the majority of the latter being private wells. over 700 000 drinking-water samples are registered in jupiter, with almost 10 million analytical measurements of water quality parameters. data sources locations of public supplies were initially extracted from jupiter in october 2013 and have been continuously updated since. missing coordinates were geocoded using address information (schullehner & hansen 2014). information on wsas is often available at municipality level. the 98 danish municipalities are required to devise water supply plans (in danish: vandforsyningsplaner), including which areas of the municipalities are supplied by which public water production facilities. some municipalities publish online geographic information system (gis) services with wsa polygons, which were downloaded. the majority of municipalities published maps, which were collected and digitised into polygons in gis. for larger water supply companies, especially around the larger cities, waterworks and/or municipalities were contacted and supplied either files with supply zone polygons or distribution networks or a description of their supply infrastructure and zones (schullehner & hansen 2014; schullehner et al. 2017). data processing in the initial version, used in a number of epidemiological studies (schullehner et al. 2018, 2019, 2020; coffman et al. 2021; thygesen et al. 2021; wodschow et al. 2021; richter et al. 2022; theisen et al. 2022), it was assumed that the wsa polygons did not change over time. the water production facility identifier (plantid, in danish: anlaegid) was spatially joined to the wsa polygons. in most cases, the water production facility is located within the wsa polygon that it supplies. in other cases, wsas were created as multipolygons including a small polygon around the plant’s address, such that plants are located within the wsa polygons they supply. for the large water supply companies of the major cities, links between plantid and the wsa identifier (wsaid) have https://doi.org/10.34194/geusb.v49.8319 http://www.geusbulletin.org schullehner 2022: geus bulletin 49. 8319. https://doi.org/10.34194/geusb.v49.8319 3 of 6 www.geusbul let in.org been added manually to the linkage table using information supplied by these companies. for historical assessments of drinking-water quality, changes in supply infrastructure can be important. therefore, the initial model was updated and has already been applied in recent epidemiological studies (stayner et al. 2021, 2022; thomsen et al. 2021; ebdrup et al. 2022). in the updated model, the static polygons – corresponding to the infrastructure at the time of data curation – were individually assessed. if more than one water production facility was located within a wsa, the wsa was split into multiple historical wsas and linked to the corresponding plantid. the start and end year of the wsa’s activity was estimated by assessing the time series of drinking-water quality samples linked to the respective plantid in jupiter. if a time series of, for example, nitrate measurements at a production facility stopped, it was assumed that this plantid closed and was merged with the wsa of the plantid with the continuing time series. this results in two polygons with each their wsa being merged into one, with a new wsaid. the start and end year of the active period of the wsas was updated: the former wsas’ active periods end in the year when the time series stopped, and the merged wsa’s start is the following year, resulting in continuity in linkage of supplied households and their water production facilities. the borders of the delineation between wsas to split were estimated by expert assessment. map layers of locations of households and private wells were used to guide the decision. in ambiguous cases, additional information was acquired from waterworks’ homepages, municipality archives and personal contact with waterworks. as geocoded residential history of the danish population is available in high quality from 1978 onwards (cirrau 2022), and drinkingwater quality monitoring achieved good coverage from the 1980s (see data description and main features); the period included in this data set is 1978–2019. software data management was done in r (versions 3.2–4.0; r core team 2020). gis tasks were done in arcmap 10 and qgis (versions 2.14–3.18; esri 2015; qgis.org 2022). data description and main features the data set presented here consists of the wsa polygons with attributes wsaid (unique key), start and end (start and end year of the wsa’s active period) as a shapefile. additionally, there is a table linking wsaid to the unique key of the water production facilities in jupiter (plantid, in danish: anlaegid). figure 1 shows the extent of the estimated wsa boundaries in 1978 and 2019. note that whilst most municipalities divide their entire area into potential wsas, some (e.g., aalborg municipality) provided the more precise physical extents, leaving areas without public supply empty. even though a household lies a: 1978 0 b: 2019 50 100 km fig. 1 extents of the water supply areas (wsas) in denmark in (a) 1978 and (b) 2019. light grey areas are outside of public water supply zones. https://doi.org/10.34194/geusb.v49.8319 http://www.geusbulletin.org http://qgis.org schullehner 2022: geus bulletin 49. 8319. https://doi.org/10.34194/geusb.v49.8319 4 of 6 www.geusbul let in.org within a wsa, it may be supplied by a private well (see later discussion on limitations). during the study period, a consolidation of the water supply structure can be observed: the number of active wsas decreased from initially 3172 in 1978 to 2602 in 2019 (fig. 2a). this corresponds to closures of water production facilities and consequent mergers with neighbouring wsas. mergers of wsas also happen without closures of production facilities, resulting in the merged wsa being supplied by more than one production facility. in 1978, 74% of wsas were supplied by only a single production facility; however, by 2019, this decreased to 66%. figure 2b shows the number of waterworks with registered drinking-water samples in jupiter. before 1990, there are reporting issues depending on the county in which the waterworks were located. after 1990, the number of waterworks with a registered drinking-water sample as an indication of being active decreases from more than 3500 to 2500. changes in the monitoring programme are also highlighted (fig. 2b). while major parameters like nitrate and calcium have been monitored consistently from the 1980s onwards, emerging quality issues such as pesticides and perand polyfluoroalkyl substances (pfas) were introduced at later stages. the arsenic drinking-water standard in denmark was lowered to 5 µg/l in 2001 (ramsay et al. 2021), resulting in increased monitoring. figure 3 shows how this data set bridges the gap when linking drinking-water quality data from jupiter to a person’s estimated exposure history in an epidemiological study. whilst jupiter, this data set, and the danish address register (sdfe 2022) are all publicly accessible data sets, getting access to the administrative and health registers requires permissions, which can be granted to researchers under specific terms complying with data protection legislation (thygesen et al. 2011). figure 3 is a minimum working example and can be extended thoroughly in both directions. linking additional information from jupiter, such as raw water quality and abstraction well geology, and additional data from administrative and health registers, such as socioeconomic factors and comorbidities, is easily accomplished using the plantid and personal identification number (pin), respectively. the strength of this data set is the establishment of the missing link between longitudinal nationwide water quality at production facility level and supply area. in combination with geocoded addresses linked to the civil registration system, links to health and administrative databases are possible. this data set takes infrastructural changes into account, by allowing 2600 2800 3000 3200 1980 1990 2000 2010 2020 w sa a 0 1000 2000 3000 1980 1990 2000 2010 2020 year pu bl ic w at er wo rk s sampled for: any parameter nitrate calcium pesticides arsenic pfasb fig. 2 annual number of active water supply areas (wsa; a) and number of public waterworks with drinking-water samples analysed for different parameters registered in jupiter (b). to account for different sampling frequencies, the annual number is presented here as the rolling count over three years. https://doi.org/10.34194/geusb.v49.8319 http://www.geusbulletin.org schullehner 2022: geus bulletin 49. 8319. https://doi.org/10.34194/geusb.v49.8319 5 of 6 www.geusbul let in.org wsa polygons and their link to production facilities to change over time. a major limitation is that this data set is not authoritative, which limits its application for other uses besides research. boundaries of wsas can be uncertain, but typically run through areas with very sparse population density, reducing the risk of misclassification. for the wsas supplied by multiple production facilities, the variation within the wsa may even change during the course of a day depending on production and consumption patterns, but this cannot be estimated with this data set alone. however, it is possible to weigh mean concentrations of water quality parameters by annual production volume registered in jupiter. when using this data set, it is, therefore, suggested to identify wsas with large variations for the specific parameter under investigation and examine these more carefully. in this data set, private wells (supplying less than ten households) are not included. however, it is possible to estimate which households are supplied by private wells, for example by proximity analysis (schullehner et al. 2017). it must, however, be noted that private well registration is not complete in jupiter. furthermore, it should be noted that raw data in jupiter often require substantial filtering, clean-up and data management before being useful for nationwide drinking-water quality mapping. the nature of this process depends on many circumstances and should be assessed individually for every waterquality parameter and study period, as, for example, data coverage, analytical methods and detection limits can vary substantially. it is, therefore, not possible here to provide a curated data set for all drinking-water quality parameter data in jupiter. research groups who made the effort to curate parameter-specific jupiter data sets are encouraged to publish these and relevant code along with their studies. supported by its wide use in epidemiological studies, it is concluded that this data set is a valuable tool for research purposes. given its limitations, it is cautioned against using this data set for administrative applications by public authorities. instead, it is encouraged that the authorities work towards establishing a database, for example, in connection with the jupiter infrastructure, where waterworks report and update officiary data. for example, by supplying spatial data with their actual supply zones or distribution systems or lists of the supplied addresses. data availability and updates the data set is available at the geus bulletin dataverse (see additional files), where future updates will be uploaded, under the creative commons licence attribution 4.0 international (cc by 4.0). it would be greatly appreciated if users of this data set report any errors or inconsistencies that they encounter. acknowledgements i would like to thank my colleagues denitza voutchkova and birgitte hansen, geological survey of denmark and greenland, for feedback and discussions about the construction of this data set, and the reviewers, whose comments improved the quality of this article. additional information funding statement this data set has been developed during several years with funding from innovation fund denmark (dnmark research alliance), the centre for register-based research at aarhus university (cirrau), the aarhus university research foundation (grant auff-e-2015-fls-8-61), pltchemanalysis sampleid drinking-water quality analyses pltchemsample sampleid plantid drinking-water sample data jupiter wsa_plant wsaid plantid start end this data set civil registration system pin residential history address_id cancer register pin diagnosis_code diagnosis_date administrative and health registers danish address register address_id x, y spatial join wsaid = 2159 fig. 3 minimum working example of using this data set for establishing the linkage between drinking-water quality in jupiter to a person’s estimated exposure history in epidemiological studies. pltchemanalysis and pltchemsample refer to the tables in jupiter containing the drinking-water plants’ chemical analyses and sample data. https://doi.org/10.34194/geusb.v49.8319 http://www.geusbulletin.org schullehner 2022: geus bulletin 49. 8319. https://doi.org/10.34194/geusb.v49.8319 6 of 6 www.geusbul let in.org the united states national institutes of health/national institute of environmental health sciences (grant r01 es027823-01a1), the karen elise jensens foundation and the novo nordisk foundation challenge programme bertha – the danish big data centre for environment and health (grant nnf17oc0027864). competing interests none. author contributions js is the sole author. additional files the complete data set is available at https://doi.org/10.22008/fk2/ i5r1ss references cirrau. 2022: data documentation – geocodes. https://cirrau.au.dk/ data-resources/data-documentation (accessed april 2022) coffman, v.r. et al. 2021: prenatal exposure to nitrate from drinking water and markers of fetal growth restriction: a population-based study of nearly one million danish-born children. environmental health perspectives 129(2), 027002. https://doi.org/10.1289/ehp7331 ebdrup, n.h. et al. 2022: nitrate in drinking water and time to 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https://doi.org/10.3390/geosciences5010045 voutchkova, d.d., schullehner, j., skaarup, c., wodschow, k., ersbøll, a.k. & hansen, b. 2021: estimating pesticides in public drinking water at the household level in denmark. geus bulletin 47, 6090. https://doi. org/10.34194/geusb.v47.6090 wodschow, k., hansen, b., schullehner, j. & ersbøll, a.k. 2018: stability of major geogenic cations in drinking water-an issue of public health importance: a danish study, 1980-2017. international journal of environmental research and public health 15(9), 1212. https://doi. org/10.3390/ijerph15061212 wodschow, k., villanueva, c.m., larsen, m.l., gislason, g., schullehner, j., hansen, b. & ersbøll, a.k. 2021: association between magnesium in drinking water and atrial fibrillation incidence: a nationwide population-based cohort study, 2002–2015. environmental health 20, 126. https://doi.org/10.1186/s12940-021-00813-z https://doi.org/10.34194/geusb.v49.8319 http://www.geusbulletin.org 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https://doi.org/10.1016/j.gexplo.2017.03.006 https://doi.org/10.1016/j.gexplo.2017.03.006 https://doi.org/10.1016/j.schres.2019.06.016 https://doi.org/10.1289/ehp6391 https://eng.sdfe.dk/products-and-services/the-danish-address-register https://eng.sdfe.dk/products-and-services/the-danish-address-register https://doi.org/10.3390/ijerph19020823 https://doi.org/10.3390/ijerph19020823 https://doi.org/10.1016/j.envint.2021.106613 https://doi.org/10.1016/j.lanepe.2021.100286 https://doi.org/10.1016/j.lanepe.2021.100286 https://doi.org/10.1016/j.envint.2022.107277 https://doi.org/10.1186/s12940-021-00805-z https://doi.org/10.1186/s12940-021-00805-z https://doi.org/10.1177/1403494811399956 https://doi.org/10.1177/1403494811399956 https://doi.org/10.1016/j.jtemb.2021.126828 https://doi.org/10.3390/geosciences5010045 https://doi.org/10.34194/geusb.v47.6090 https://doi.org/10.34194/geusb.v47.6090 https://doi.org/10.3390/ijerph15061212 https://doi.org/10.3390/ijerph15061212 https://doi.org/10.1186/s12940-021-00813-z danish water supply areas and their links to water production facilities: an open-access data set abstract tabular abstract introduction data collection setting data sources data processing software data description and main features data availability and updates acknowledgements additional information funding statement competing interests author contributions additional files references figures fig. 1 extents of the water supply areas (wsas) in denmark in (a) 1978 and (b) 2019. light grey areas are outside of public water supply zones. fig. 2 annual number of active water supply areas (wsa; a) and number of public waterworks with drinking-water samples analysed for different parameters registered in jupiter (b). to account for different sampling frequencies, the annual number is presented here as the rolling count over three years. fig. 3 minimum working example of using this data set for establishing the linkage between drinking-water quality in jupiter to a person’s estimated exposure history in epidemiological studies. pltchemanalysis and pltchemsample refer to the tables in jupiter containing the drinking-water plants’ chemical analyses and sample data. bojesen-koefoed et al. 2023: geus bulletin 55. 8350. https://doi.org/10.34194/geusb.v55.8350 1 of 8 preface the rødryggen-1 and brorson halvø-1 fully cored boreholes (upper jurassic – lower cretaceous), wollaston forland, north-east greenland – an introduction jørgen a. bojesen-koefoed1* , peter alsen2 , morten bjerager3 , jussi hovikoski3,4 , jon r. ineson3 , peter n. johannessen5, mette olivarius2 , stefan piasecki3,6,7 , henrik vosgerau3 1department for mapping and mineral resources, geological survey of denmark and greenland (geus), copenhagen, denmark; 2department for geo-energy and storage, geological survey of denmark and greenland (geus), copenhagen, denmark; 3department for geophysics and sedimentary basins, geological survey of denmark and greenland (geus), copenhagen, denmark; 4geological survey of finland (gtk), espoo, finland; 5geological survey of denmark and greenland (geus), now retired; 6globe institute, university of copenhagen, copenhagen, denmark; 7retired abstract two fully cored boreholes, the rødryggen-1 and the brorson halvø-1, were drilled in wollaston forland, north-east greenland, in 2009 and 2010, respectively. the objective was to test the stratigraphic development of the upper jurassic – lower cretaceous mud-dominated succession in two different settings within the same fault block of a developing half-graben: centrally (rødryggen-1 borehole) and near the uplifted crest of the rotating fault block (brorson halvø-1 borehole). the drilled deposits are equivalent to the principal petroleum source-rock sequence of the petroliferous basins of north-west europe, siberia, and basins off eastern canada and provide a new record of an important phase of marine deoxygenation in the proto-north atlantic region. *correspondence: jbk@geus.dk received: 26 may 2023 revised: 27 jun 2023 accepted: 27 jun 2023 published: 21 dec 2023 keywords: stratigraphic boreholes, technical data, sedimentology, biostratigraphy, geochemistry abbreviations: geus: geological survey of denmark and greenland td: total depth gr: gamma ray b.rfl.: below reference level geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: karen dybkjær (geus, denmark) reviewed by: not peer reviewed. funding: see page 7 competing interests: none. additional files: none. 1. introduction in 2007–2008, the geological survey of denmark and greenland (geus) initiated a major collaborative project with a consortium of companies of the international petroleum industry with the aim of addressing aspects of the geology of north-east and north greenland, which were generally recognised as being insufficiently studied. one such aspect was the elucidation of the imperfectly known stratigraphy, sedimentary and lithologic development, and petroleum generation potential of the upper jurassic – lower cretaceous succession, equivalent to the well-known petroleum source-rock succession of the north sea graben system, that is, the kimmeridge clay sensu lato. in north-east greenland, the equivalent succession is widely distributed, and a considerable number of outcrop samples as well as some samples from cored shallow boreholes had been analysed previously, generally showing surprisingly poor petroleum generation potential. this was suspected to be the result of a combination of poor representation in the sample set of the more prolific intervals and adverse effects of weathering of outcrop samples. hence, a drilling program was designed to obtain full core representation of the entire stratigraphic succession ranging from the oxfordian to the ryazanian. three fully cored boreholes were planned, the first of which, the blokelv-1 borehole, was drilled in jameson land in 2008. blokelv-1 penetrated 233.8 m of oxfordian–volgian deposits, https://doi.org/10.34194/geusb.v55.8350 https://orcid.org/0000-0001-5647-2769 https://orcid.org/0000-0001-6218-9054 https://orcid.org/0000-0003-3180-8857 https://orcid.org/0000-0001-6330-8713 https://orcid.org/0000-0003-0017-3705 https://orcid.org/0000-0003-3853-7543 https://orcid.org/0000-0002-7846-859x https://orcid.org/0000-0002-7582-5360 mailto:jbk@geus.dk bojesen-koefoed et al. 2023: geus bulletin 55. 8350. https://doi.org/10.34194/geusb.v55.8350 2 of 8 geusbulletin.org intruded by a few basaltic sills, as reported in a series of papers published in a volume edited by ineson & bojesen-koefoed (2018). two additional boreholes, both situated in wollaston forland, the rødryggen-1 and brorson halvø-1 boreholes, were drilled in 2009 and 2010, respectively (fig. 1). these boreholes aimed at testing that part of the succession not drilled by the blokelv-1 borehole whilst attaining a reasonable stratigraphic overlap. the two boreholes were planned to be drilled centrally within, and near the uplifted crest of, a rather narrow half-graben, which evolved during the time of deposition, thus allowing the assessment of the effects of synsedimentary tectonic movements on the nature of the deposits. the main results of the drilling are reported in a series of papers dealing with stratigraphy (alsen et al. 2023, this volume): sedimentology and basin evolution (hovikoski et al. 2023a, this volume), mineralogy and diagenesis (olivarius et al. 2023, this volume) and organic geochemistry and petroleum potential (bojesen-koefoed et al. 2023, this volume). the aim of this paper is to introduce the location, drilled successions, drilling procedure and technical specifications of the rødryggen-1 and brorson halvø-1 boreholes. 18°w 16°w 76°n 72°n 22°w 20°w ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ g re en la nd traill ø hochstetter forland store koldewey wollaston forland jameson land 100 km 26°w 22°w24°w28°w 26°w 24°w " 74°45'n 74°50'n 74°55'n 74°40ʹn 74°35ʹn 74°20ʹn 74°15ʹn 74°25ʹn 74°30ʹn 19°30'w 19°0'w 20°0ʹw20°30ʹw brorson halvø-1 albrecht bugt daneborg p. f. k .f . h .f . rødryggen-1 w o l l a s t o n f o r l a n d k u h n ø young sund clavering ø s abine ø falske bugt major dykes/sills faults plateau lavas ice lakes rivers quaternary stratumbjerg and fosdalen fms palnatokes bjerg fm lindemans bugt fm bernbjerg fm pelion, jakobsstigen and payer dal fms permian pre-caledonian basement 10 km palaeogene borehole d .f . cretaceous jurassic triassic permian fault inferred deep-seated fault ■ ■ fig. 1 location maps. overview map (left; modified from surlyk et al. 2023). positions of the rødryggen-1 and brorson halvø-1 boreholes in wollaston forland are shown on a simplified geological map (right; modified from surlyk et al. 2021). k.f.: kuhn fault; p.f.: permpas fault; h.f.: hühnerbjerg fault; d.f.: dombjerg fault. the dombjerg fault was the main fault to control the position of the coastline during the late jurassic. the permpas–hühnerbjerg block(s) was bounded by the kuppel and hühnerbjerg faults during the late jurassic. https://doi.org/10.34194/geusb.v54.8350 http://www.geusbulletin.org/ bojesen-koefoed et al. 2023: geus bulletin 55. 8350. https://doi.org/10.34194/geusb.v55.8350 3 of 8 geusbulletin.org 2. drilling and samples 2.1 rødryggen-1 borehole drilling and technical specifications of the rødryggen-1 fully cored borehole are summarised in table 1. the rødryggen-1 borehole was drilled from 27 july to 16 august 2009 at the rødryggen locality, central wollaston forland. the drilling rig and the drill camp were mobilised by helicopter (sling and cabin load) from the military station at daneborg approximately 30 km to the ssw of the drill site (fig. 1). rødryggen (“red ridge”) is a prominent ns-oriented, elongated hill, separating two broad depressions: storsletten to the west and sumpdalen to the east. the drill-site was selected during the 2008 field season by henrik nøhr-hansen and jørgen a. bojesen-koefoed. the well location was chosen with the aim of testing the upper jurassic – lower cretaceous succession centrally in the westward-tilted permpas– hühnerbjerg fault block. the rig was placed on outcrops of the bright yellow-weathering mudstones of the albrechts bugt member of the palnatokes bjerg formation (figs 2 and 3). debris from the overlying rødryggen member of the palnatokes bjerg formation unit imparts a conspicuous red colour to the rødryggen hill and thus forms the basis for its name. the rig and the drilling team spent 22 days on the location, with only 11 effective days of drilling. the marked discrepancy between time on location and time spent on active drilling was caused by prolonged periods of very bad weather with strong winds and large volumes of snow and rain, which tended to liquefy the upper decimeters of the very muddy ground, thus creating massive problems for camping and for the stability of the rig. the drilled succession is summarised in table 2. the borehole penetrated 24.4 m of the albrechts bugt member of the palnatokes bjerg formation (light grey, yellow-weathering mudstones), followed by 72.6 m of the lindemans bugt formation (dark grey/black laminated mudstones), a succession that was later defined as a new member, the storsletten member (alsen et al. 2023, this volume). beneath the lindemans bugt formation, the borehole encountered 137.5 m of the bernbjerg formation (dark grey mudstones with occasional thin sandstone stringers and carbonate-filled thin fractures) before reaching total depth (td) at 234.5 m below reference level (b.rfl.; top of casing). for details of the sedimentology, the mineralogy and diagenesis, and the organic geochemistry and petroleum potential see hovikoski et al. (2023a, this volume) olivarius et al. (2023, this volume) and bojesen-koefoed et al. (2023, this volume), respectively. a total gamma-ray (gr) wireline log was collected in the borehole over the interval 0–209 m b.rfl., later supplemented by spectral-gr and bulk-density scanning of the cores over the interval 192.56 m to 234.54 m. table 1 drilling and technical data on the rødryggen-1 and brorson halvø-1 boreholes. specifications rødryggen-1 brorson halvø-1 borehole number ggu 517001 ggu 517003 borehole name rødryggen-1 brorson halvø-1 area wollaston forland, north-east greenland wollaston forland, north-east greenland operator geus geus drilling operator geus geus altitude 110 m above mean sea level 101 m above mean sea level coordinates (wgs 84) 74°32.561’n; 19°50.924’w 74°35.227’n; 19°34.327’w utm zone 27w 27w easting e465752 e457646 northing n8272914 n8278142 drill rig sandvik de 130 sandvik de 130 casing diameter (o/i) 64/57 mm 64/57 mm casing depth 15.8 m 35.0 m borehole diameter 56 mm 56 mm core diameter 42 mm 42 mm reference level top of casing top of casing total depth 234.5 m 225.6 m core recovery 99% 99% status abandoned open hole, top of casing closed with a steel cap. abandoned open hole. hole collapsed. drilling crew arrival (daneborg) 24 july 2009 n.a. mobilisation to drill site 27 july – 29 july 2009 30 july – 1 august 2010 preparation of camp 27 july – 2 august 2009 2–3 august2010 spud 3 august 2009 3 august 2010 drilling completed 14 august 2009 12 august 2010 demobilisation 16 august 2009 14 august 2010 effective drilling 11 days 10 days total days at drill location 22 days 14 days https://doi.org/10.34194/geusb.v54.8350 http://www.geusbulletin.org/ bojesen-koefoed et al. 2023: geus bulletin 55. 8350. https://doi.org/10.34194/geusb.v55.8350 4 of 8 geusbulletin.org fig. 2 the rødryggen-1 drill-site and camp, viewed approximately towards the west with the broad plain of storsletten as backdrop (photo: annette ryge). the rig was placed on the brightly coloured, yellow-weathering mudstones of the albrechts bugt member (palnatokes bjerg formation). the overlying rødryggen member (palnatokes bjerg formation) in the foreground consists of purple to brick-red mudstones. debris from this unit imparts the slopes of the landscape with a reddish colour that can be seen from afar, and from which is derived the location’s name rødryggen, meaning “the red ridge.” table 2 stratigraphic units encountered by the rødryggen-1 and brorson halvø-1 boreholes. lithostratigraphy rødryggen-1 brorson halvø-1 stratumbjerg formation n.d. 8.7 m (0–8.7 m) palnatokes bjerg formation, rødryggen member n.d. 21.3 m (8.7–30 m) palnatokes bjerg formation, albrects bugt member 24.4 m (0–24.4 m) 7.5 m (30–37.5 m) lindemans bugt formation, storsletten member 72.6 m (24.4–97 m) 8.0 m (37.5–45.5 m) bernbjerg formation 137.5 m (97–234.5 m) 180.1 m (45.5–225.6 m) thickness in m (intervals in m b.rfl.). n.d.: not drilled. storsletten mb. (lindemans bugt fm.) albrechts bugt mb. (palnatokes bjerg fm.) rødryggen mb. (palnatokes bjerg fm.) stratumbjerg fm. fig. 3 section at rødryggen viewed approximately from the north in 2008 when the drill-site was selected. arrow indicates the approximate position of the borehole (photo: jørgen a. bojesen-koefoed). https://doi.org/10.34194/geusb.v54.8350 http://www.geusbulletin.org/ bojesen-koefoed et al. 2023: geus bulletin 55. 8350. https://doi.org/10.34194/geusb.v55.8350 5 of 8 geusbulletin.org the wireline log and the scanned gr core logs were then merged into one single total gr log. the core recovery was 99%, with losses of core primarily occurring in the uppermost part of the succession, which was affected by surface weathering. the core quality was mostly excellent, with significant fracturing only found in intervals scattered throughout the succession, whilst coherent core in general predominated (fig. 4). a total of 70 whole core samples for gas analysis, each c. 10 cm in length, were collected on site at approximately 3 m intervals and stored in airtight containers immediately after the cores were extruded from the core barrel. a large number of samples for different types of analyses were subsequently collected from the core. most of these analyses are reported in papers by alsen et al. (2023, this volume), hovikoski et al. (2023a, this volume), olivarius et al. (2023, this volume) and bojesen-koefoed et al. (2023, this volume). additional data are reported by hovikoski et al. (2023b), whereas a minor volume of miscellaneous data that have not found a place in the publications listed above are reported in internal geus reports. 2.2 brorson halvø-1 borehole drilling and technical specifications of the brorson halvø-1 fully cored borehole are summarised in table 1. the brorson halvø-1 borehole was drilled from 30 july to 14 august 2010 at brorson halvø (brorson peninsula), northeastern wollaston forland. the brorson halvø-1 borehole was the second to be drilled during the 2010 field season, and the drilling rig and camp were mobilised from the first drill site on the island of store koldewey, approximately 200 km north of the brorson halvø drill-site (fig. 1). the rig frame was transported by helicopter in one piece as sling load directly from the store koldewey drill site to the brorson halvø drill site. the drilling team, the drill camp and other equipment were moved either directly by helicopter or via danmarkshavn by twin otter plane to daneborg and from there by helicopter to the brorson halvø-1 drill-site. the original drill site was selected during the 2008 field season by henrik nøhr-hansen and jørgen a. bojesen-koefoed, later redefined by morten bjerager and michael b.w. fyhn in 2009, and eventually, for practical reasons, moved another few 100 m by stefan piasecki and jørgen a. bojesen-koefoed in 2010 to secure a stable substratum for the rig. a consequence of this last relocation of the drill site was the drilling of part of the stratumbjerg formation and the entire rødryggen member of the palnatokes bjerg formation, which was initially not planned. the well location was selected with the aim of testing the upper jurassic – lower cretaceous succession near the uplifted crest of the westward-tilted permpas-hühnerbjerg fault block. the rig was placed on outcrops of fig. 4 example of core quality and recovery, rødryggen-1 borehole, 59.81 – 66.16 m, lindemans bugt formation (photo: john boserup). https://doi.org/10.34194/geusb.v54.8350 http://www.geusbulletin.org/ bojesen-koefoed et al. 2023: geus bulletin 55. 8350. https://doi.org/10.34194/geusb.v55.8350 6 of 8 geusbulletin.org the dark-grey mudstones of the stratumbjerg formation (fig. 5). the rig and the drilling team spent 14 days on the location, with 10 effective days of drilling. the drilled succession is summarised in table 2. the borehole penetrated 8.7 m of the stratumbjerg formation (dark-grey mudstones), followed by 21.3 m of the rødryggen member (purple to brick-red mudstones) and 7.5 m of the albrechts bugt member (light grey, yellow-weathering mudstones) of the palnatokes bjerg formation. beneath the palnatokes bjerg formation, the borehole encountered 8.0 m of the lindemans bugt formation (dark grey/ black laminated mudstones), later referred to the new storsletten member (alsen et al. 2023, this volume), and eventually, 180.1 m of the bernbjerg formation (dark grey mudstones with occasional thin sandstone stringers and carbonate-filled, thin fractures). td was reached fig. 5 the brorson halvø-1 drill site viewed approximately towards the ene. the drill-rig stands on the dark grey mudstones of the stratumbjerg formation with the well-defined contact to the bright red rødryggen member clearly visible below. downhill, the contact between the rødryggen member and the underlying albrechts bugt member (yellow) stands out. the contact to the storsletten member (grey), in turn, underlying the albrechts bugt member appears somewhat gradational, but in a clean section, it is very well defined (see fig. 6; photo: annette ryge). fig. 6 example of core quality and recovery, brorson halvø-1 borehole, 34.44–41.24 m, albrechts bugt member of the palnatokes bjerg formation (light grey) and the sharp boundary with the underlying storsletten member of the lindemans bugt formation (dark grey; photo: john boserup). https://doi.org/10.34194/geusb.v54.8350 http://www.geusbulletin.org/ bojesen-koefoed et al. 2023: geus bulletin 55. 8350. https://doi.org/10.34194/geusb.v55.8350 7 of 8 geusbulletin.org at 225.6m b.rfl. (top of casing). drilling was terminated due to notable water production from the borehole at td. for details of the sedimentology, the mineralogy and diagenesis, and the organic geochemistry and petroleum potential, see hovikoski et al. (2023a, this volume), olivarius et al. (2023, this volume) and bojesen-koefoed et al. (2023, this volume), respectively. a total gr wireline log was collected in the borehole over the interval 0–210 m b.rfl., later supplemented by spectral-gr and bulk-density scanning of the cores over the interval 193 m – 225.6 m. the wireline log and the scanned gr core logs were then merged into one single total gr log. the core recovery was 99%, with significant losses of recovery mainly in the uppermost part of the succession, which was affected by surface weathering. the core quality was mostly excellent, with significant fracturing only found in intervals scattered throughout the succession, whilst coherent core predominated (fig. 6). a total of 34 whole-core samples for gas analysis, each c. 10 cm in length, were collected on site and stored in airtight containers immediately after the cores were extruded from the core barrel. a large number of samples for different types of analyses were subsequently collected from the core. most of these analyses are reported in papers by alsen et al. (2023, this volume), hovikoski et al. (2023a, this volume), olivarius et al. (2023, this volume) and bojesen-koefoed et al. (2023, this volume). additional data are reported by hovikoski et al. (2023b), whereas a minor volume of miscellaneous data that have not found a place in the publications listed above are reported in internal geus reports. 2.3 perspectives the primary data from the three fully cored, complementary boreholes blokelv-1, rødryggen-1 and brorson halvø-1 are presented in ineson & bojesen-koefoed (2018) and this volume. together these cored sections represent a unique geological archive from a high latitude setting, recording a key period when super-regional deoxygenation prevailed in the marine waters of the proto-north atlantic region. though the significance of this archive may have waned commercially, given the imperative to adopt carbon-free energy sources, its continued importance lies in understanding the climatic, tectonic and oceanographic factors involved in watermass stratification, deoxygenation and carbon burial, as examined recently by hovikoski et al. (2023b). it is predicted that this integrated core record will be the source of much multidisciplinary research in future years. acknowledgements the drilling teams including john boserup, anders clausen, annette ryge, peter turner, lars (lasse) thomsson, andreas hjort frandsen and anders pilgaard are thanked for their hard work and stamina, and for maintaining good spirits, particularly during the very challenging 2009 field season. the pilots finn rusanes and göran lindmark (both air greenland) are thanked for diligent sling-work under very difficult conditions. twin otter plane support by nordlandair (akureyri, iceland) is gratefully acknowledged. jesper weiss andersen and thea weiss andersen (in 2009 only) did an excellent job running the logistics of the 2009 and 2010 expeditions. additional information funding statement the drilling campaign was funded by geus in collaboration with a large number of companies of the international petroleum industry, which for contractual reasons cannot be named. author contributions jabk: conceptualisation; writing – original draft; funding acquisition; project administration; investigation; writing – review & editing. pal: conceptualisation; funding acquisition; investigation; writing – review & editing. mbj: investigation; writing – review & editing. jhov: investigation; writing – review & editing. ji: conceptualisation; writing – original draft: investigation; writing – review & editing. pj: investigation. mol: investigation; writing – review & editing. sp: conceptualisation; investigation; writing – review & editing. hv: investigation; writing – review & editing. competing interests none declared additional files none provided references alsen, p., piasecki, s., nøhr-hansen, h., pauly, s., sheldon, e. & hovikoski, j. 2023: stratigraphy of the upper jurassic to lowermost cretaceous in the rødryggen-1 and brorson halvø-1 boreholes, wollaston forland, north-east greenland. geus bulletin 55, 8342 (this volume). https:// doi.org/https://doi.org/10.34194/geusb.v55.8342 bojesen-koefoed, j.a., alsen, p., bjerager, m., hovikoski, j., johannessen, p., nøhr-hansen, h., petersen, h.i., piasecki, s. & vosgerau, h. 2023: organic geochemistry of an upper jurassic – lower cretaceous mudstone succession in a narrow graben setting, wollaston forland basin, north-east greenland. geus bulletin 55, 8320 (this volume). https:// doi.org/https://doi.org/10.34194/geusb.v55.8320 hovikoski, j., ineson, j.r, olivarius, m., bojesen-koefoed, j.a, piasecki, s. & alsen, p. 2023: upper jurassic – lower cretaceous of eastern wollaston forland, north-east greenland: a distal marine record of an evolving rift. geus bulletin 55, 8349 (this volume). https://doi.org/ https://doi.org/10.34194/geusb.v55.8349 hovikoski, j. et al. 2023b: late jurassic – early cretaceous marine deoxygenation in ne greenland. journal of the geological society 180 (3), jgs2022–058. https://doi.org/10.1144/jgs2022-058 ineson, j.r. & bojesen-koefoed, j.a. (eds). 2018: petroleum geology of the upper jurassic – lower cretaceous of east and north-east greenland: blokelv-1 borehole, jameson land basin. geological survey of denmark and greenland bulletin 42, 168pp. https://doi.org/10.34194/geusb.v42 olivarius, m., kazerouni, a.m., weibel, r., kokfelt, t.f. & hovikoski, j. 2023: mudstone diagenesis and sandstone provenance in an upper jurassic – lower cretaceous evolving half-graben system, wollaston forland, north-east greenland. geus bulletin 55, 8309 (this volume). https://doi.org/https://doi.org/10.34194/geusb.v55.8309 surlyk, f. 1978: submarine fan sedimentation along fault scarps on tilted fault blocks (jurassic–cretaceous boundary, east greenland). bulletin grønlands geologiske undersøgelse 128, 117 pp. https://doi. org/10.34194/bullggu.v128.6670 surlyk, f. 2003: the jurassic of east greenland: a sedimentary record of thermal subsidence, onset and culmination of rifting. in: ineson, j.r. https://doi.org/10.34194/geusb.v54.8350 http://www.geusbulletin.org/ https://doi.org/https https://doi.org/https http://doi.org/10.34194/geusb.v55.8342 https://doi.org/https https://doi.org/https http://doi.org/10.34194/geusb.v55.8320 https://doi.org/https https://doi.org/https http://doi.org/10.34194/geusb.v55.8349 https://doi.org/10.1144/jgs2022-058 https://doi.org/10.34194/geusb.v42 https://doi.org/https http://doi.org/10.34194/geusb.v55.8309 https://doi.org/10.34194/bullggu.v128.6670 https://doi.org/10.34194/bullggu.v128.6670 bojesen-koefoed et al. 2023: geus bulletin 55. 8350. https://doi.org/10.34194/geusb.v55.8350 8 of 8 geusbulletin.org & surlyk, f. (eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 659–722. https://doi. org/10.34194/geusb.v1.4674 surlyk, f. et al. 2021: jurassic stratigraphy of east greenland. geus bulletin 46, 6521. https://doi.org/10.34194/geusb.v46.6521 surlyk, f., alsen, p., hovikovski, j. & piasecki, s. 2023: uplift, deflation and marine onlap of a jurassic rift dome, illustrated by a backstepping middle – upper jurassic shelf-to-slope succession, geographical society ø, east greenland. terra nova 00, 1–8. https://doi.org/10.1111/ ter.12673 https://doi.org/10.34194/geusb.v54.8350 http://www.geusbulletin.org/ https://doi.org/10.34194/geusb.v1.4674 https://doi.org/10.34194/geusb.v1.4674 https://doi.org/10.34194/geusb.v46.6521 https://doi.org/10.1111/ ter.12673 https://doi.org/10.1111/ ter.12673 the rødryggen-1 and brorson halvø-1 fully cored boreholes (upper jurassic – lower cretaceous), wollaston forland, north-east greenland – an introduction 1. introduction 2. drilling and samples 2.1 rødryggen-1 borehole 2.2 brorson halvø-1 borehole 2.3 perspectives acknowledgements additional information funding statement author contributions competing interests additional files references figures fig. 1 location maps. overview map (left; modified from surlyk et al. 2023). positions of the rødryggen-1 and brorson halvø-1 boreholes in wollaston forland are shown on a simplified geological map (right; modified from surlyk et al. 2021). k.f.: kuhn fault; p.f.: permpas fault; h.f.: hühnerbjerg fault; d.f.: dombjerg fault. the dombjerg fault was the main fault to control the position of the coastline during the late jurassic. the permpas–hühnerbjerg fig. 2 the rødryggen-1 drill-site and camp, viewed approximately towards the west with the broad plain of storsletten as backdrop (photo: annette ryge). the rig was placed on the brightly coloured, yellow-weathering mudstones of the albrechts bugt member (palnatokes bjerg formation). the overlying rødryggen member (palnatokes bjerg formation) in the foreground consists of purple to brick-red mudstones. debris from this unit imparts the slopes of the landscape with a reddish colour that can be seen from afar, and from which is derived the location’s name rødryggen, meaning “the red ridge.” fig. 3 section at rødryggen viewed approximately from the north in 2008 when the drill-site was selected. arrow indicates the approximate position of the borehole (photo: jørgen a. bojesen-koefoed). fig. 4 example of core quality and recovery, rødryggen-1 borehole, 59.81 – 66.16 m, lindemans bugt formation (photo: john boserup). fig. 5 the brorson halvø-1 drill site viewed approximately towards the ene. the drill-rig stands on the dark grey mudstones of the stratumbjerg formation with the well-defined contact to the bright red rødryggen member clearly visible below. downhill, the contact between the rødryggen member and the underlying albrechts bugt member (yellow) stands out. the contact to the storsletten member (grey), in turn, underlying the albrechts bugt member appears somewhat gradational, but in a clean section, it is very well defined (see fig. 6; photo: annette ryge). fig. 6 example of core quality and recovery, brorson halvø-1 borehole, 34.44–41.24 m, albrechts bugt member of the palnatokes bjerg formation (light grey) and the sharp boundary with the underlying storsletten member of the lindemans bugt formation (dark grey; photo: john boserup). tables table 1 drilling and technical data on the rødryggen-1 and brorson halvø-1 boreholes. table 2 stratigraphic units encountered by the rødryggen-1 and brorson halvø-1 boreholes. bojesen-koefoed et al. 2023: geus bulletin 55. 8320. https://doi.org/10.34194/geusb.v55.8320 1 of 24 research article organic geochemistry of an upper jurassic – lower cretaceous mudstone succession in a narrow graben setting, wollaston forland basin, north-east greenland jørgen a. bojesen-koefoed1* , peter alsen2 , morten bjerager3 , jussi hovikoski3,4 , peter n. johannessen5, henrik nøhr-hansen2 , henrik i. petersen2 , stefan piasecki3,6,7 , henrik vosgerau3 1department for mapping and mineral resources, geological survey of denmark and greenland (geus), copenhagen, denmark; 2department for geo-energy and storage, geological survey of denmark and greenland (geus), copenhagen, denmark; 3department for geophysics and sedimentary basins, geological survey of denmark and greenland (geus), copenhagen, denmark; 4geological survey of finland (gtk), espoo, finland; 5geological survey of denmark and greenland (geus), now retired; 6globe institute, university of copenhagen, copenhagen, denmark; 7retired abstract the oxfordian–ryazanian was a period of widespread deposition of marine organic-rich mudstones in basins formed during the early phases of the rifting that heralded the formation of the present-day north atlantic. occasionally, uninterrupted deposition prevailed for 20 million years or more. today, mudstones of this time interval are found on the shelves bordering the north atlantic and adjacent areas from siberia to the netherlands. here, we report data on two fully cored boreholes from wollaston forland (north-east greenland, approx. 74° n), which represent an uninterrupted succession from the upper kimmeridgian to the hauterivian. the boreholes record basin development at two different positions within an evolving halfgraben, located at the margin of the main rift, and thus partially detached from it. although the overall depositional environment remained an oxygen-restricted deep-shelf setting, rifting-related changes can be followed through the succession. the kimmeridgian was a period of eustatic highstand and records the incipient rifting with a transgressive trend straddling the transition to the lower volgian by a gradual change from deposits with high levels of total organic carbon (toc) and kerogen rich in allochthonous organic matter to deposits with lower toc and a higher proportion of autochthonous organic matter. this is followed by a slight regressive trend with lower toc and increased proportions of allochthonous organic matter until rifting culminated in the middle volgian–ryazanian, indicated by increasing autochthonous organic matter and higher toc, which prevailed until basin ventilation occurred towards the end of the ryazanian. the properties of the reactive kerogen fraction remained rather stable irrespective of toc, underlining the effect of terrigenous matter input for toc. these variations are also captured by biological markers and stable carbon isotopes. the deposits are very similar to equivalent successions elsewhere in the proto-north atlantic region, albeit the proportion of terrigenous kerogen is greater. *correspondence: jbk@geus.dk received: 20 apr 2022 revised: 06 jan 2023 accepted: 10 jan 2023 published: 21 dec 2023 keywords: north-east greenland, organic geochemistry, paleogeography, source rock, upper jurassic abbreviations: b. rfl.: below reference level gc: gas chromatography geus: geological survey of denmark and greenland hi: hydrogen index hilive: average hi of the live kerogen fraction ms: mass spectrometry mmboe: million barrels of oil equivalent nso: nitrogen, sulphur and oxygen pi: production index sra: source rock analyzer tc: total carbon td: total depth tmax: temperature at maximum rate of pyrolysate generation during rock-eval or sra analysis (°c) toc: total organic carbon ts: total sulphur ueg: ultimate expulsion gas ueo: ultimate expulsion oil uep: ultimate expulsion potential geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: jon r. ineson (geus, denmark) reviewed by: erdem idiz (university of oxford, uk), iain scotchman (scotchman geochemistry services ltd, uk) funding: see page 22 competing interests: see page 22 additional files: none provided 1. introduction marine shales of oxfordian–ryazanian age constitute the most important source rocks for petroleum in the prospective basins of the north atlantic region, which include the basins of the greater north sea area, the barents shelf, and the basins west of ireland and the shetland islands as well as the basins of western siberia. on the western side of the atlantic this also includes https://doi.org/10.34194/geusb.v55.8320 https://orcid.org/0000-0001-5647-2769 https://orcid.org/0000-0001-6218-9054 https://orcid.org/0000-0003-3180-8857 https://orcid.org/0000-0001-6330-8713 https://orcid.org/0000-0002-9291-8104 https://orcid.org/0000-0001-6606-7062 https://orcid.org/0000-0002-7846-859x https://orcid.org/0000-0002-7582-5360 mailto:jbk@geus.dk https://creativecommons.org/licenses/by/4.0/ bojesen-koefoed et al. 2023: geus bulletin 55. 8320. https://doi.org/10.34194/geusb.v55.8320 2 of 24 geusbulletin.org the jeanne d’arc and flemish pass basins off eastern canada, and probably untested basins off east and north-east greenland. these deposits have been extensively studied (e.g. von der dick et al. 1989; miller 1990; chakhmakhchev et al. 1994; klemme 1994; telnæs et al. 1994; fowler & mcalpine 1995; isaksen & ledje 2001; ineson et al. 2003; justwan & dahl 2005; justwan et al. 2005, 2006a,b; petersen et al. 2010; scotchman et al. 2016). age equivalent deposits crop out onshore north-east greenland. however, irrespective of their importance for the general understanding of the most important petroleum system of northwest europe, published in-depth studies of the nature and petroleum potential of the equivalent northeast greenland succession are scarce. exceptions include papers by requejo et al. (1989), christiansen et al. (1992), strogen et al. (2005) and bojesen-koefoed et al. (2018). over the years 2008–2010, the geological survey of denmark and greenland (geus) drilled three fully cored boreholes to depths of more than 200 m to penetrate the upper jurassic – lower cretaceous mudstone succession in east and north-east greenland. the successions drilled are partially time equivalents of the kimmeridge clay formation of the wessex basin, uk, as well as of the upper jurassic – lower cretaceous petroleum source-rock successions of the north sea and north atlantic basins, where they are known under a variety of different local names, see for instance ineson et al. (2003). the kimmeridge clay formation has been penetrated by cored boreholes close to its type section in dorset (morgans-bell et al. 2001). however, the east and north-east greenland boreholes offer an opportunity to study nearly the full oxfordian to ryazanian succession in an area remote from other studied outcrops and wells. the first of these boreholes to be drilled was the blokelv-1 in jameson land, which covers the succession from the oxfordian to the lower volgian (see ineson & bojesen-koefoed 2018). the second and third of the planned boreholes, the rødryggen-1 and brorson halvø-1, respectively, were drilled in northern wollaston forland, north-east greenland in 2009 and 2010, respectively. the brorson halvø-1 drill site is situated approximately 10 km north-east of the rødryggen-1 drillsite, near the uplifted eastern crest of the fault block defined by the permpas fault to the west and the hühnerbjerg fault to the east (fig. 1; surlyk 1978). the rødryggen-1 borehole is located near the centre of the same block. the main target of the drilling in both cores was the same kimmeridgian–ryazanian black mudstone succession with the primary objective being to delineate the lateral development in facies of the upper jurassic – lower cretaceous mudstone succession in an evolving half graben system in the wollaston forland area. the objective of this paper is to present an overview of the organic geochemistry of the upper jurassic – lower cretaceous mudstone succession in the wollaston forland basin, based on new evidence from the rødryggen-1 and brorson halvø1 boreholes (fig. 1). details of the drilling, sedimentology and stratigraphy of the rødryggen-1 and brorson halvø1 boreholes can be found in bojesen koefoed et al. (2023, this volume), hovikoski et al. (2023b, this volume) and alsen et al. (2023, this volume). 2. geological setting and stratigraphy of the drilled succession the jurassic–cretaceous wollaston forland basin developed in response to rifting in the proto-north atlantic to the east of the basin. the wollaston forland basin is overall a system of westerly tilted fault blocks or half-grabens bounded to the west by the dombjerg fault and to the east by the hühernbjerg fault (figs 1, 2; surlyk, 1978, 2003; fyhn et al. 2021a). the basin is internally segmented into several subbasins defined by roughly north–south trending normal faults, which were active during various phases of the basin development (fyhn et  al. 2021a,b; hovikoski et al. 2023a,b and references therein). the two drill-cores penetrate the kimmeridgian – lower barremian succession, which includes four lithostratigraphic units (fig. 3): (1) the bernbjerg formation, (2) the lindemans bugt formation, storsletten member (alsen et al. 2023, this volume), (3) the palnatokes bjerg formation (the albrechts bugt member and the rødryggen member), and (4) the stratumbjerg formation (bjerager et al. 2020; surlyk et al. 2021). the upper oxfordian – lower volgian bernbjerg formation is up to 500–600 m thick and represents a tectonically affected muddy shelf succession that crops out from store koldewey in the north to traill ø (ø meaning island) in the south (surlyk & clemmensen 1983; surlyk 2003; surlyk et  al. 2021). a major rift episode starting in the volgian terminated the regionally continuous shelf accumulation and resulted in accelerated tilted fault-block development and basin segmentation that lasted until the barremian in the western part of the area (surlyk 1978, 1984, 1990, 2003; piasecki et al. 2020). as a result, the basin geometry changed into a series of narrow, 10–30 km wide, s–n-oriented basins that were strongly westwards tilted. in the most proximal fault block, the volgian–ryazanian syn-rift interval is characterised by major conglomeratic submarine fandelta systems (lindemans bugt formation; rigi member; surlyk 1978; henstra et  al. 2016), which graded into heterolithic and mud-dominated deposits in more distal areas (the laugeites ravine, niesen and storsletten members). the areal distribution of the lindemans bugt formation is limited to north-eastern clavering ø, north-western wollaston forland, eastern th. thomsen land (named  for the ethnographer thomas thomsen, known as th. thomsen (higgins 2010)) and south-west https://doi.org/10.34194/geusb.v55.8320 http://www.geusbulletin.org/ bojesen-koefoed et al. 2023: geus bulletin 55. 8320. https://doi.org/10.34194/geusb.v55.8320 3 of 24 geusbulletin.org fig. 1 location maps. overview map (left). positions of the rødryggen-1 and brorson halvø-1 boreholes on wollaston forland are shown on a simplified geological map (right). k.f.: kuhn fault; p.f.: permpas fault; h.f.: hühnerbjerg fault; d.f.: dombjerg fault. the dombjerg fault was the main fault to control the position of the coastline during the late jurassic. the permpas–hühnerbjerg block(s) was bounded by the kuppel and hühnerbjerg faults, which probably represented the main controlling faults in the block that is studied here, during the late jurassic. reproduced from bojesen-koefoed et al. (2023, this volume). fig. 2 conceptual, approximately sw–ne-oriented cross-section of clavering ø – wollaston forland, showing multiple westward-tilted fault blocks. modified from birkelund & perch-nielsen (1976) and vischer (1943). 18°w 16°w 76°n 72°n 22°w 20°w ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ g re en la nd traill ø hochstetter forland store koldewey wollaston forland jameson land 100 km 26°w 22°w24°w28°w 26°w 24°w " 74°45'n 74°50'n 74°55'n 74°40ʹn 74°35ʹn 74°20ʹn 74°15ʹn 74°25ʹn 74°30ʹn 19°30'w 19°0'w 20°0ʹw20°30ʹw brorson halvø-1 albrecht bugt daneborg p. f. k .f . h .f . rødryggen-1 w o l l a s t o n f o r l a n d k u h n ø young sund clavering ø s abine ø falske bugt major dykes/sills faults plateau lavas ice lakes rivers quaternary stratumbjerg and fosdalen fms palnatokes bjerg fm lindemans bugt fm bernbjerg fm pelion, jakobsstigen and payer dal fms permian pre-caledonian basement 10 km palaeogene borehole d .f . cretaceous jurassic triassic permian fault inferred deep-seated fault ■ ■ pro�le rudis bugt louis elv clavering ø skillegletscher young sund wollaston forland blæsedalen m 0 4000 plateau basalt cretaceous jurassic triassic upper permian carboniferous and lower permian caledonian basement sea https://doi.org/10.34194/geusb.v55.8320 http://www.geusbulletin.org/ bojesen-koefoed et al. 2023: geus bulletin 55. 8320. https://doi.org/10.34194/geusb.v55.8320 4 of 24 geusbulletin.org kuhn ø (surlyk 1978; surlyk et al. 2021). the formation is wedge-shaped in a west–east direction and is estimated to reach a maximum thickness of 2 km. the valanginian stage is characterised by waning rift activity in the study area (late syn-rift) and transgression, whereas rifting continued in the axial areas to the east (surlyk 1978, 1984, 2003; surlyk & korstgård 2013; hovikoski et al. 2018). the palnatokes bjerg formation was deposited during this period. the formation crops out in wollaston forland, kuhn ø, hochstetter forland and traill ø (surlyk 1978; surlyk et al. 2021). like the lindemans bugt formation, the formation shows variable thickness in west–east transect and reaches a maximum thickness of 600 m. it includes the coarse-grained gravity flow deposits of the young sund and falskebugt members, and the fossiliferous fine-grained albrechts bugt and rødryggen members. the cored interval penetrates both fine-grained members and the recent biostratigraphic data suggest a valanginan to hauterivian age for these deposits (alsen & mutterlose 2009; pauly et al. 2012; möller et al. 2015). in the brorson halvø core, the palnatokes bjerg formation is gradationally overlain by a thin interval of upper hauterivian sub-storm-wave-base bioturbated mudstones of the stratumbjerg formation (bjerager et  al. 2020). this formation crops out from traill ø in the south to store koldewey in the north and reaches  its maximum thickness of 270 m in the brorson halvø area. 3. samples and methods samples used for the present study include material from the fully cored rødryggen-1 and brorson halvø-1 boreholes. the rødryggen-1 borehole reached a total depth (td) of 234.66 m below terrain, with a core recovery of 99%. the lithostratigraphic units encountered include  the bernbjerg formation (kimmeridgian – lower volgian), the lindemans bugt formation (lower volgian – upper ryazanian) and the palnatokes bjerg formation, albrechts bugt member (upper ryazanian – upper valanginian; alsen et al. 2023, this volume). the succession penetrated by fig. 3 stratigraphic models of the wollaston forland – kuhn ø area. left: the original stratigraphy by surlyk (2003). right: the revised stratigraphy after drilling of the rødryggen-1 and brorson halvø-1 boreholes and incorporating changes introduced by surlyk et al. (2021). the revised succession is much more complete and richer in mudstone than suggested by the original. east greenl. tect. strat. seq.: east greenland tectono-stratigraphic sequence. u: upper. m: middle. l: lower. abbreviations for geology are as follows: str: stratumbjerg formation. r: rødryggen member. f: falske bugt member. pa: palnatokes bjerg formation (young sund member). al: albrechts bugt member. n: niesen member. li: lindemans bugt formation. li (s): lindemans bugt formation (storsletten member). ri: rigi member. l: laugeites ravine member. b: bernbjerg formation. ug: ugpik ravine member. j: jakobsstigen formation. pay: payer dal formation. pe: pelion formation. mu: muslingebjerg formation. ba: bastians dal formation. black dashes indicate uncertain contact. black vertical lines indicate boreholes, r-1: rødryggen-1; bh-1: brorson halvø-1. chronostratigraphywollaston forland – kuhn ø east greenl. tect. strat. seq. r falpa li ri n b j pe ug pay mu ba onlaps crystalline basementonlaps upper permian w e s n s n 2.6 2.5 2.4 2.3 2.2 2.1 hauterivian valanginian ryazanian volgian berriasian tithonian kimmeridgian oxfordian callovian alluvial/delta plain – paralic, sand-dominated coal shallow marine sandstones shelf transition – sandstones, mudstones, heteroliths o�shore/basinal mudstones deep marine sandstones deep marine conglomerates calcareous sandy marine mudstones red marine mudstones hiatus/condensed bathonian bajocian aalenian c re ta ce ou s ju ra ss ic lo w er u pp er m id dl e u m l u l u m l u m l u m l u l u m l l chronostratigraphywollaston forland – kuhn ø r ? al str bh-1 r-1pa no data li li (s)ri n b j pe ug pay mu ba onlaps crystalline basementonlaps upper permian w e barremian hauterivian valanginian ryazanian volgian berriasian tithonian kimmeridgian oxfordian callovian bathonian bajocian aalenian c re ta ce ou s ju ra ss ic lo w er u pp er m id dl e u m u l l u m l u m l u m l u l u m l l ? https://doi.org/10.34194/geusb.v55.8320 http://www.geusbulletin.org/ bojesen-koefoed et al. 2023: geus bulletin 55. 8320. https://doi.org/10.34194/geusb.v55.8320 5 of 24 geusbulletin.org the borehole is stratigraphically complete and does not include any notable hiatuses. a total of 258 samples were subjected to total carbon (tc), total sulphur (ts), total organic carbon (toc) and rock-eval type screening analyses, and subsets of these samples were selected for further analyses such as vitrinite reflectance analysis (12 samples), biological marker analysis (24 samples) and stable carbon isotopic analysis (20 samples). the brorson halvø-1 borehole reached a td of 225.88 m below terrain, with a core recovery of 99%. the  lithostratigraphic units encountered include the bernbjerg formation (kimmeridgian – lower volgian), the lindemans bugt formation (middle volgian), the palnatokes bjerg formation including the albrechts bugt member (upper ryazanian – upper valanginian), and rødryggen member (hauterivian) and the stratumbjerg formation (barremian; alsen et al. 2023, this volume). the succession penetrated by the borehole is stratigraphically incomplete and includes notable hiatuses with significant portions of the upper part of the lower volgian and the lower part of the middle volgian being absent. a significant portion of the upper part of the middle volgian, the entire upper volgian and lower ryazanian successions are also missing (fig. 3). a total of 232 samples were subjected to tc/ts/toc-rockeval type screening analyses, and subsets of these samples were selected for further analyses, such as vitrinite reflectance analysis (10 samples), biological marker analysis (18 samples) and stable carbon isotopic analysis (10 samples). in addition to the borehole samples, a set of outcrop samples collected in the immediate vicinity of the rødryggen-1 drill site were subjected to tc/ts/ toc-rock-eval type screening analyses. the sample set represents dense sampling of a c. 30 m thick profile of the uppermost part of the storsletten member (lindemans bugt formation), extending stratigraphically downwards from the well-defined boundary between the storsletten member of the lindemans bugt formation and the overlying albrechts bugt member. the boundary thus serves as a datum for the sampling, which can also be recognised in the rødryggen-1 core. the outcrop samples were collected from regular outcrops exposed by digging away the cover of loose shale debris. analytical procedures, summarised here, are detailed in full in bojesen-koefoed et al. (2018). tc (wt%), toc (wt%) and ts (wt%) were determined by combustion in a leco cs-200 induction furnace. petroleum potential was determined by rock-eval type pyrolysis using a source rock analyzer (sra) instrument, manufactured by humble instruments and services and calibrated against the ifp160000 standard. particulate blocks for reflected light microscopy were prepared and measured for vitrinite reflectance according to international standards (taylor et al. 1998). several samples were also qualitatively inspected in reflected white light and fluorescence-inducing blue light. solvent extraction (samples powdered to <250 µm) was  carried out with methanol/dichloromethane 7:93 vol./vol. as solvent using a soxtec™. asphaltenes were precipitated by addition of 40-fold excess n-pentane. maltene fractions were separated into saturated, aromatic and nso fractions (nso: compounds containing nitrogen, sulphur and oxygen and other heteroatoms) by medium-pressure liquid chromatography (radke et al. 1980). gas chromatography of saturated extract fractions was carried out using a shimadzu gc-2010 instrument. gas chromatography (gc) – mass spectrometry (ms) was carried out using an agilent 6890n gas chromatograph connected to a waters (micromass) quattro micro gc tandem quadrupole-hexapole-quadrupole mass spectrometer. 4. results 4.1 thermal maturity the thermal maturity of the successions penetrated by the rødryggen-1 and brorson halvø-1 boreholes was assessed using a combination of several independent parameters including the temperature at the maximum rate of pyrolysate generation (°c) tmax and production index (pi) derived from rock-eval type pyrolysis, vitrinite reflectance (%ro), and sterane isomerization ratios (tables 1, 2, 4, 5). fig.  4, upper panel, shows maturity data for both boreholes drilled versus depth. fig.  4, lower panel, likewise shows maturity data versus depth but includes data on the stratigraphic breakdown and the presence of hiatuses in the brorson halvø-1 succession documented by alsen et  al. (2023, this volume). hence, the brorson halvø-1 data are shifted to accommodate two major hiatuses, assuming the thickness of missing sections equals the thickness of corresponding sections in the rødryggen-1 borehole. despite this very simplistic approach, a surprisingly good match is obtained when considering detailed variations, see fig. 5, which has been prepared in a similar way. in both boreholes, homohopane isomerization ratios have reached equilibrium distribution, and thus carry no information on the maturity gradient, but demonstrate that the burial temperature has exceeded that required for the equilibrium reaction to have been completed. the rødryggen-1 borehole shows clearly increasing trends with depth in all parameters (fig. 4). tmax increases from c. 420°c in the ryazanian succession to 430°c in the kimmeridgian succession at the base of the borehole. pi shows a parallel increase from c. 0.04 to 0.08, whereas the vitrinite reflectance increases from c. 0.48% ro to 0.61% ro, and sterane 20s/(20s+20r) isomerization ratio https://doi.org/10.34194/geusb.v55.8320 http://www.geusbulletin.org/ bojesen-koefoed et al. 2023: geus bulletin 55. 8320. https://doi.org/10.34194/geusb.v55.8320 6 of 24 geusbulletin.org goes from 0.30 to 0.49. only the sterane αββ/(ααα+αββ) isomerization ratio shows a slightly more irregular depth trend, but on average, this ratio increases from c. 0.29 to 0.36 over the succession penetrated by the rødryggen-1 borehole. combined, the maturity parameters agree in suggesting that the succession is thermally immature fig. 4 thermal maturity parameters versus drilled depth below reference level (b. rfl.) for the rødryggen-1 (red symbols/lines) and brorson halvø-1 (blue symbols/lines) boreholes (upper panels) and the rødryggen-1 (red symbols/lines) and brorson halvø-1 (blue symbols/lines) boreholes taking into account stratigraphic information (lower panels; alsen et al. 2023, this volume). two hiatuses in the succession penetrated by the brorson halvø-1 borehole have been compensated for by assuming that the thickness of the missing section equals the equivalent section in the rødryggen-1 borehole, which shows no hiatus. hence, only samples of the rødryggen-1 borehole show true drilled depths, while samples below the hiatus in the brorson halvø-1 section have been shifted to greater depths and may even appear deeper than the total depth (td) of the brorson halvø-1 borehole. tmax: temperature of maximum rate of generation of pyrolysate during rock-eval type pyrolysis. production index: s1/(s1+s2) from rock-eval type pyrolysis. ro (%): vitrinite reflectance. s29 20s/(20s+20r): c29 sterane 20αααs/(20αααs+20αααr) isomer ratio. s29 αββ/(αββ+ααα): c29 sterane 20αββ/(20ααα+20αββ) isomer ratio. total dia/regular: total diasteranes to total regular steranes ratio. 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 0 0.1 0.2 0.3 0.4 0.5 production index 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 0.2 0.3 0.4 0.5 0.6 s29 20s/(20s+20r) 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 0.2 0.4 0.6 0.8 s29 αββ/(αββ+ααα) 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 0 1 2 3 4 5 total dia/regular 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 m et re s b. r � . 400 420 440 460 480 tmax (°c) 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 0.4 0.6 0.8 1 ro (%) 300 290 280 270 260 250 240 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 0 0.1 0.2 0.3 0.4 0.5 production index 300 290 280 270 260 250 240 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 0.2 0.3 0.4 0.5 0.6 s29 20s/(20s+20r) 300 290 280 270 260 250 240 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 0.2 0.3 0.4 0.5 0.6 0.7 300 290 280 270 260 250 240 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 0 1 2 3 4 5 total dia/regular 300 290 280 270 260 250 240 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 m et re s b. r � . 400 420 440 460 480 tmax (°c) 300 290 280 270 260 250 240 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 0.4 0.6 0.8 1 ro (%) s29 αββ/(αββ+ααα) https://doi.org/10.34194/geusb.v55.8320 http://www.geusbulletin.org/ bojesen-koefoed et al. 2023: geus bulletin 55. 8320. https://doi.org/10.34194/geusb.v55.8320 7 of 24 geusbulletin.org with respect to petroleum generation but is approaching oil-window maturity at the base of the borehole. the existence of a clear maturity gradient over such a limited depth interval is unexpected, but a similar feature is found in the blokelv-1 borehole in jameson land (bojesen-koefoed et al. 2018), and in the brorson halvø-1 borehole, situated approximately 10  km north-east of the  rødryggen-1 drill site. over the succession penetrated by the brorson halvø-1 borehole, tmax increases from c. 435°c to 450°c with parallel increases in pi from c. 0.09 to 0.19, and in vitrinite reflectance from 0.54% ro to 0.85% ro. sterane 20s/(20s+20r) isomerization ratios have reached equilibrium distribution in the entire succession, whereas the sterane αββ/(ααα+αββ) isomer ratio increases from c. 0.55 to 0.62 with depth. combined, the maturity parameters suggest that the upper part of the succession just enters the oil-generative window, which extends further towards the base of the borehole. comparing the two panels of fig. 4, the hiatus has little detectable effect on the maturity profile of the brorson halvø-1 borehole. despite the similar depth trends, which suggest a regionally high geothermal gradient, there is a notable difference in the level of thermal maturity observed in the two boreholes. 4.2 petroleum potential and organic facies variations the succession penetrated by the rødryggen-1 borehole is stratigraphically complete with no notable hiatus. the kimmeridgian and lower volgian sections show a steady up-section decrease in toc, from close to 6 wt% to c. 3 wt% at the transition to the middle volgian (fig. 5; table 1). ts remains largely constant at c. 2.5 wt% over the same interval, whereas the s2-parameter displays some variation, which is shown by the derived hydrogen index (hi) that shows initial values close to 200 and a gradual upwards increase culminating locally with a hi of c. 300, some 35 m below the transition to the lower volgian. the section above shows a decreasing trend with a minimum hi of c. 200 in the lowermost part of the lower volgian section, followed by a notable increase towards the base of the middle volgian to hi close to 350. average hi remains largely constant at 300–350 in the middle volgian and the lower portion of the ryazanian sections, after which it essentially drops to zero. however, in general terms, the drilled succession shows only limited variation in kerogen type as defined by screening data, i.e. a gas–oil-prone kerogen type ii/iii (fig. 6). superimposed on the general trends described here are a number of subordinate trends. both general and subordinate trends can often be tied to transgressive-regressive cycles indicated by detailed sedimentological analysis and various chemical proxies, and which in turn are linked to the tectonic evolution of the area as described by hovikoski et al. (2023a,b). the stratigraphically partly overlapping succession penetrated by the brorson halvø-1 borehole shows very similar detailed variations to the corresponding succession in the rødryggen-1 borehole, except for the effects of increased thermal maturity, as shown by higher s1-values and lower s2-values and hi in the brorson halvø-1 borehole (figs 5, 7; table 2). moreover, due to the presence of hiatuses, certain intervals are not present. the uninterrupted sedimentary record demonstrated by the succession drilled by the rødryggen-1 borehole and its low level of thermal maturity allow a more detailed assessment of the petroleum potential of the individual chronostratigraphic intervals. using the method of dahl et al. (2004), the average proportions of inert/dead carbon and the average hi of the live kerogen fraction (hilive) have been assessed (fig. 8; table 3). the results show that the kimmeridgian, lower volgian, middle volgian and upper volgian successions all include similar proportions of dead carbon, yielding values roughly in the range 1.5–2.0 wt%. conversely, hilive tends to increase steadily up-section to culminate at a value of 610 in the upper volgian section. the ryazanian succession shows very low proportions of inert/dead carbon and slightly lower hilive than the upper volgian section, but the data set is limited and includes non-source deposits. several biological marker parameters and stable carbon isotope ratio data (δ13c) show trends that follow those outlined for the screening data. representative chromatograms and ion fragmentograms for both boreholes are shown in figs 9 and 10. key biological markers and δ13c data are listed in tables 4 and 5, and selected biological marker parameters and δ13c data are shown versus depth for both boreholes in fig. 11. with some scatter, the pristane/phytane ratio shows a slight, but clear upwards-decreasing trend over the entire drilled succession, with no clear differences between the two boreholes. the isohopane ratio (nytoft 2011) shows a much better defined, but very similar trend, i.e. a clear decrease from the base of the succession until the base of the middle volgian succession from where the ratio seems to become stable. the homohopane ratio (peters et al. 2005 and references therein) shows a clear increase upwards through the succession until the middle volgian where the ratio becomes invariant. the boreholes show identical trends, but the brorson halvø-1 data are shifted in the order of 5 percentage points towards higher values. the gammacerane index in the rødryggen-1 borehole shows upwards-increase into the lower volgian, followed by a decrease continuing into the middle volgian and a sharp increase and stabilisation in https://doi.org/10.34194/geusb.v55.8320 http://www.geusbulletin.org/ bojesen-koefoed et al. 2023: geus bulletin 55. 8320. https://doi.org/10.34194/geusb.v55.8320 8 of 24 geusbulletin.org the middle volgian through the ryazanian section. in the brorson halvø-1 borehole, no clear trends are observed. the proportion of c30 desmethyl steranes shows a steadily upwards-increasing trend throughout the entire penetrated successions with no obvious differences observed between the two boreholes. the δ13c data shows a slight upwards-decrease in the kimmeridgian section, which becomes more pronounced upwards through the lower volgian section followed by stabilisation or even a slight increase in the middle fig. 5 organic geochemical screening parameters versus drilled depth below reference level (b. rfl.) for the rødryggen-1 (red symbols/lines) and brorson halvø-1 (blue symbols/lines) boreholes (upper panels) and the rødryggen-1 (red symbols/lines) and brorson halvø-1 (blue symbols/lines) boreholes taking into account stratigraphic information (lower panels; alsen et al. 2023, this volume). two hiatuses in the succession penetrated by the brorson halvø-1 borehole have been compensated for by assuming that the thickness of the missing section equals the equivalent section in the rødryggen-1 borehole, which shows no hiatus. hence, only samples of the rødryggen-1 borehole show true drilled depths, while samples below the hiatuses in the brorson halvø-1 section have been shifted to greater depths and may even appear deeper than the total depth (td) of the brorson halvø-1 borehole. stratigraphic ages as defined by the rødryggen-1 borehole. 300 290 280 270 260 250 240 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 0 2 4 6 8 10 toc (wt%) 300 290 280 270 260 250 240 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 0 1 2 3 4 s1 (kg hc / ton rock) 300 290 280 270 260 250 240 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 0 4 8 12 16 20 s2 (kg hc / ton rock) 300 290 280 270 260 250 240 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 0 200 400 600 hydrogen index 300 290 280 270 260 250 240 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 m et re s b. r � . 0 4 8 12 tc (wt%) 300 290 280 270 260 250 240 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 0 4 8 12 16 ts (wt%) kim merid gian l. v olgian m. v olgian u. volgian ryaza nian 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 0 2 4 6 8 10 toc (wt%) 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 0 1 2 3 4 s1 (kg hc / ton rock) 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 0 4 8 12 16 20 s2 (kg hc / ton rock) 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 0 200 400 600 hydrogen index 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 m et re s b. r � . 0 4 8 12 tc (wt%) 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 0 4 8 12 16 ts (wt%) https://doi.org/10.34194/geusb.v55.8320 http://www.geusbulletin.org/ bojesen-koefoed et al. 2023: geus bulletin 55. 8320. https://doi.org/10.34194/geusb.v55.8320 9 of 24 geusbulletin.org table 1 summary of organic geochemical data per stratigraphic interval of the rødryggen-1 borehole ryazanian upper volgian middle volgian lower volgian kimmeridgian (n = 23) (n = 9) (n = 55) (n = 55) (n = 97) tc minimum 0.81 3.55 2.49 2.80 4.28 maximum 5.03 5.36 5.60 10.19 10.66 mean 3.14 4.39 4.07 4.55 6.00 ts minimum 0.02 2.52 1.58 0.73 0.70 maximum 6.49 8.80 12.05 4.21 5.91 mean 3.10 5.04 3.75 2.21 2.72 toc minimum 0.10 3.07 2.11 1.48 1.30 maximum 4.77 4.68 5.37 5.61 7.22 mean 2.37 3.98 3.72 3.91 5.36 tmax minimum 419 419 420 421 424 maximum 428 425 430 428 432 mean 421 422 424 425 428 s1 minimum 0.01 0.28 0.25 0.22 0.26 maximum 0.80 0.60 1.05 1.20 1.58 mean 0.46 0.47 0.64 0.60 1.01 s2 minimum 0.33 8.93 5.39 2.36 2.77 maximum 19.46 15.41 19.83 14.98 17.99 mean 13.58 12.89 12.61 9.13 12.76 hi minimum 39 271 214 136 137 maximum 445 363 431 342 308 mean 355 324 335 231 238 pi minimum 0.03 0.03 0.03 0.04 0.05 maximum 0.04 0.04 0.07 0.09 0.13 mean 0.03 0.04 0.05 0.06 0.07 tc: total carbon (wt%). ts: total sulphur (wt%). toc: total carbon (wt%). tmax: tmax (°c) from rock-eval pyrolysis, s1: free hydrocarbons (mg/g) from rock-eval pyrolysis. s2: pyrolytic hydrocarbons (mg/g) from rock-eval pyrolysis. hi: hydrogen index (100 × s2/toc). pi: production index (s1/(s1+s2)). table 2 summary of organic geochemical data per stratigraphic interval of the brorson halvø-1 borehole ryazanian middle volgian lower volgian kimmeridgian (n = 3) (n = 8) (n = 125) (n = 62) tc minimum 0.71 2.62 2.86 4.58 maximum 1.28 3.93 3.91 10.95 mean 0.99 3.54 3.48 6.35 ts minimum 3.76 2.44 2.44 1.03 maximum 5.79 9.63 9.63 6.72 mean 4.57 5.46 4.74 3.07 toc minimum 0.04 2.23 1.76 1.64 maximum 0.10 3.79 6.58 8.07 mean 0.06 3.08 4.05 5.74 tmax minimum 406 433 436 443 maximum 435 438 438 452 mean 423 436 437 447 s1 minimum 0.00 0.76 0.76 0.48 maximum 0.00 1.67 1.67 2.91 mean 0.00 1.19 1.25 2.23 s2 minimum 0.00 6.52 6.52 1.87 maximum 0.00 10.94 10.94 15.77 mean 0.00 9.07 8.61 10.64 hi minimum 0 250 250 114 maximum 0 324 325 220 mean 0 294 303 184 pi minimum n.a. 0.07 0.10 0.13 maximum n.a. 0.20 0.20 0.36 mean n.a. 0.12 0.13 0.18 tc: total carbon (wt%). ts: total sulphur (wt%). toc: total carbon (wt%). tmax: tmax (°c) from rock-eval pyrolysis. s1: free hydrocarbons (mg/g) from rock-eval pyrolysis. s2: pyrolytic hydrocarbons (mg/g) from rock-eval pyrolysis. hi: hydrogen index (100 × s2/toc). pi: production index (s1/(s1+s2)). https://doi.org/10.34194/geusb.v55.8320 http://www.geusbulletin.org/ bojesen-koefoed et al. 2023: geus bulletin 55. 8320. https://doi.org/10.34194/geusb.v55.8320 10 of 24 geusbulletin.org fig. 6 standard plots of organic geochemical screening data from the rødryggen-1 borehole. 380 390 400 410 420 430 440 450 460 470 480 490 500 510 520 tmax (°c) 0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1 pr od uc ti on in de x (s 1/ (s 1+ s2 )) rødryggen-1 borehole ryazanian upper volgian middle volgian lower volgian kimmeridgian immature oil condensate, wet gas dry gas staining or contamination low maturity conversion 400 450 500 550 tmax (°c) 0 100 200 300 400 500 600 700 800 900 1000 h yd ro ge n in de x 0,1 1 10 100 toc (%) 0,1 1 10 100 s2 (m g h c /g ro ck ) excellent excellent good g ood fair poor poor https://doi.org/10.34194/geusb.v55.8320 http://www.geusbulletin.org/ bojesen-koefoed et al. 2023: geus bulletin 55. 8320. https://doi.org/10.34194/geusb.v55.8320 11 of 24 geusbulletin.org fig. 7 standard plots of organic geochemical screening data from the brorson halvø-1 borehole. 380 390 400 410 420 430 440 450 460 470 480 490 500 510 520 tmax (°c) 0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1 pr od uc ti on in de x (s 1/ (s 1+ s2 )) brorson halvø-1 borehole middle volgian lower volgian kimmeridgian immature oil condensate, wet gas dry gas staining or contamination low maturity conversion 400 450 500 550 tmax (°c) 0 100 200 300 400 500 600 700 800 900 1000 h yd ro ge n in de x 0.1 1 10 100 toc (%) 0.1 1 10 100 s2 (m g h c /g ro ck ) excellent excellent good g ood fair poor poor https://doi.org/10.34194/geusb.v55.8320 http://www.geusbulletin.org/ bojesen-koefoed et al. 2023: geus bulletin 55. 8320. https://doi.org/10.34194/geusb.v55.8320 12 of 24 geusbulletin.org volgian through ryazanian section. overall, the δ13c trend is mirrored in the homohopane ratio, the isohopane ratio and the proportion of c30 desmethyl steranes. moreover, δ13c and the pristane/phytane ratio is reasonably well correlated, with increasing isotopic depletion and decreasing average pristane/phytane ratio up-section (fig. 12). superimposed on the general trends described here, minor trends can be observed in several data sets, which combined with detailed sedimentology and main and trace element data are reported by hovikoski et al. (2023a). 4.3 outcrop versus borehole samples: weathering effects the proximity of the sampled outcrop profile to the rødryggen-1 drill site allows a close comparison of borehole versus outcrop data and thus assessment of weathering effects. key data for outcrop and correlative borehole samples are shown in fig. 13. the contents of toc, and hence roughly the concentration of organic matter, do not show conspicuous differences between outcrop and borehole samples, but the quality of the organic matter, using hi as a proxy, is remarkably different. with only one exception, represented by a low-value spike on the drill-core sample data curve, the hi values are consistently and significantly lower in outcrop samples compared to drill-core samples. on average, borehole data show hi values (average hi 343) that are a factor of three to four greater than the values yielded in corresponding outcrop samples (average hi 102). ts concentrations in outcrop samples (average ts = 2.4%), which due to the clayey nature of the sediments can be safely assumed to primarily represent pyritic sulphur, are approximately half of the values observed in fresh drill-core samples (average ts = 4.3%). 5. discussion 5.1 thermal maturity the presence of clear and consistent gradients in several independent thermal maturity indicators over such limited depth intervals as represented by the drilled successions of the rødryggen-1 and brorson halvø-1 boreholes is remarkable, but a similar trend was observed in the blokelv-1 borehole in jameson land (bojesen-koefoed et al. 2018). the trend can be explained by considering the processes of thermal maturation and the amount of uplift and erosion that has taken place in the wollaston forland region. based on apatite fission track data, bonow & japsen (2021) estimate that 2–3 km of sediments and volcanic deposits have been removed by erosion, which with a slightly higher than normal geothermal gradient suggests that the succession was near to or within the ‘oil window’ during the time of maximum burial. assuming a higher-than-normal geothermal gradient seems reasonable considering rifting and magmatic activity in the area. the processes associated with thermal maturation and the conversion of kerogen into petroleum are not linear but include several ‘thresholds’, known as ‘coalification jumps’ in coal-petrographic nomenclature (taylor et al. 1998). such coalification jumps are defined by rapid changes in the rates of the maturation processes. incipient petroleum generation or the start of the oil window coincides with the first coalification jump at which point several processes related to petroleum generation accelerate. this causes the non-linearity of maturation profiles commonly seen in exploration wells entering or penetrating the oil-generative window. the thermal maturity of the rødryggen-1 and brorson halvø-1 successions includes the transition into or remains within, respectively, the oil-generative window. due to rapid uplift and erosion, which removed the overlying succession, petroleum generation ceased while leaving the unusual maturity gradient. the pronounced difference in thermal maturity between the two boreholes cannot readily be explained, but it may be speculated if the brorson halvø area was more affected by magmatic intrusions than the central parts of the basin where the rødryggen-1 borehole is situated. basaltic lavas and magmatic intrusions of various sizes crop out in the brorson halvø area whereas none are known in the immediate vicinity of the rødryggen-1 borehole. the effects of magmatic intrusions and hydrothermal waters on encasing sedimentary rocks are well known (e.g. searl 1994; bojesen-koefoed et al. 2018, 2020). 5.2 petroleum potential and organic facies variations the variations in organic facies reflect large-scale variations in depositional environment linked to relative sea-level changes, which in turn can be related to phases of the rifting in the region and tectonostratigraphic development. the overall trends are cursorily discussed further here, whereas a detailed account can be found in hovikoski et  al. (2023a,b), who also describe notable differences in the detailed sedimentology of the successions drilled by the two boreholes. however, since these have little impact on the general organic geochemical character of the deposits, they are not considered here. the kimmeridgian succession in general represents an oxygen-restricted depositional environment, receiving considerable amounts of allochthonous/terrigenous organic matter. towards the end of the kimmeridgian, https://doi.org/10.34194/geusb.v55.8320 http://www.geusbulletin.org/ bojesen-koefoed et al. 2023: geus bulletin 55. 8320. https://doi.org/10.34194/geusb.v55.8320 13 of 24 geusbulletin.org fig. 8 assessment of hilive (i.e. hi of the actively petroleum-generating part of the kerogen) and dead carbon (i.e. the carbon fraction that is inert with respect to petroleum generation) per stratigraphic unit of the rødryggen-1 borehole, following the procedure of dahl et al. (2004). data are listed in table 3. s2: pyrolytic hydrocarbons (mg/g) from rock-eval pyrolysis. toc: total organic carbon. 0 4 8 12 16 20 s2 (mg hc/g rock) 0 2 4 6 8 to c (% ) ryazanian: toc = (0.216*s2) + 0.60 0 4 8 12 16 s2 (mg hc/g rock) 0 2 4 6 8 to c (% ) upper volgian: toc = (0.164*s2) + 1.87 0 4 8 12 16 20 s2 (mg hc/g rock) 0 2 4 6 8 to c (% ) middle volgian: toc = (0.179*s2) + 1.45 0 4 8 12 16 20 s2 (mg hc/g rock) 0 2 4 6 8 to c (% ) kimmeridgian: toc = (0.301*s2) + 1.42 0 4 8 12 16 s2 (mg hc/g rock) 0 2 4 6 8 to c (% ) lower volgian: toc = (0.261*s2) + 1.52 https://doi.org/10.34194/geusb.v55.8320 http://www.geusbulletin.org/ bojesen-koefoed et al. 2023: geus bulletin 55. 8320. https://doi.org/10.34194/geusb.v55.8320 14 of 24 geusbulletin.org table 3 average contents of inert carbon and hydrogen index values of the reactive kerogen fraction per stratigraphic interval of the rødryggen-1 borehole. calculated using the method of dahl et al. (2004) interval inert c (%) live kerogen hi ryazanian 0.60 463 late volgian 1.87 610 middle volgian 1.45 559 lower volgian 1.52 383 kimmeridgian 1.42 332 a gradual change occurs, pointing to increasing stagnation and reduced input of allochthonous/terrigenous organic matter and increasing proportions of autochthonous/marine organic matter in the sediments, corresponding to an overall transgression, probably eustatic in nature compounded by the effects of initial rifting (hovikoski et  al. 2023a). this is clearly demonstrated in the sediments by the decreasing levels of toc, coinciding with steadily increasing values of hi. the transgressive trend is also manifest in decreasing pristane/ phytane and isohopane ratios, increasing proportions of marine c30 desmethyl steranes and slightly decreasing δ13c (figs 5, 11). after a period of stabilisation in the latest part of the kimmeridgian, the lower volgian succession records renewed transgression and further water column stagnation, extending well into the lower part of the middle volgian. this is shown in the deposits by stable or slightly decreasing toc paralleled by increasing hi, further decreases in pristane/phytane and isohopane ratios, increasing homohopane ratio and proportion of marine c30 desmethyl steranes plus further carbon isotopic depletion leading to decreasing δ13c (figs 5, 11). overall, the succession from the lower part of the middle volgian through the ryazanian seems to represent a fairly stable highstand with sediments dominated by autochthonous organic matter with a background contribution of allochthonous terrigenous organics. minor variations may perhaps be attributed to local development of the rift. hence, most parameters remain stable or show only weak trends. the general trends in relative sea level largely, albeit not fully, conform to the findings of sneider et al. (1995) and surlyk (2003), albeit with a shift in timing suggesting somewhat younger ages of the events recorded. the kimmeridgian through ryazanian succession, however, overall records a development including a gradual decrease in terrigenous input, increasing stagnation and marine organic input. this is illustrated in fig. 12, where δ13c and the pristane/phytane ratio both decrease upwards through the section. this is also depicted as a gradual transition from ‘deltaic’ sedimentation in the kimmeridgian towards increasingly marine sedimentation in the volgian–ryazanian. this development determines the petroleum-generation potential of the resulting deposits. the entire succession shows high potential for petroleum generation, developing from predominantly gas–oil-prone kerogen in the kimmeridgian towards an increasingly oil-prone kerogen up-section (see also section 5.4 kinextm modelling). this is also evident from the calculation of the average characteristics for partial sections defined by chronostratigraphic breakdown (fig. 8; table 3). the average hi shows a steady increase from the kimmeridgian through the upper volgian succession, and at the same time the proportion of inert carbon remains rather stable. the data on the ryazanian are not fully representative for the petroleum potential since the data set also includes samples from the non-source section deposited after ventilation of the basin in the later part of the ryazanian. 5.3 outcrop versus borehole samples: weathering effects compared to correlative stratigraphic intervals in the greater north atlantic area, outcrop samples of the upper jurassic – lower cretaceous succession in east and north-east greenland often show surprisingly low petroleum potential (requejo et  al. 1989; christiansen et  al. 1992; strogen et  al. 2005; bojesen-koefoed et  al. 2018). the exact reason for this is not fully understood but based on the data reported in this paper as well as the observations of bojesen-koefoed et  al. (2018) an important cause may be the very large concentrations of finely disseminated pyrite present in the shales. weathering of pyrite will ultimately generate sulphurous and sulphuric acids, which, by their oxidising nature, may attack kerogen and reduce its petroleum potential. in addition, decomposition through microbial sulphate reduction is conceivable since several strains of sulphate-reducing bacteria seem to be facultative anaerobes, and are thus able to survive or perhaps even proliferate in oxic environments (sass & cypionka 2007). in the present case, the true hi may be reduced by a factor of 3 or more due to weathering (fig. 13). hence, although the underlying reason is not fully clear, it can be concluded that in the present case, outcrop samples are unsuitable for assessing the true potential for petroleum generation of the deposits. a common sign of weathering of pyrite is encrustations of greenish-yellow jarosite on outcrop faces, and this occurs frequently in the succession studied here. jarosite is a potassium-iron sulphate-mineral (kfe3(so4)2(oh)6), characteristically formed by pyrite weathering. several published studies report similar effects but with toc levels often much more reduced than those observed here (leythaeuser 1973; clayton & swetland 1978; raiswell & berner 1986; https://doi.org/10.34194/geusb.v55.8320 http://www.geusbulletin.org/ bojesen-koefoed et al. 2023: geus bulletin 55. 8320. https://doi.org/10.34194/geusb.v55.8320 15 of 24 geusbulletin.org fig. 9 biological marker data from the rødryggen-1 borehole. characteristic fingerprints of samples representing the upper (sample #17895, 37.88 m, upper volgian) and lower (sample #18086, 216.55 m, kimmeridgian) parts of the drilled succession. gc-ms-ms parent–daughter traces for steranes and hopanes represent the sum of five and nine parent–daughter transitions, respectively. time si gn al time si gn al time si gn al time si gn al time si gn al time si gn al time si gn al time si gn al pris ta ne pris ta ne ph yt an eph yt an e nc₁₅ nc 15 nc 20 nc₂₀ nc₂₅ nc 25 nc 30 nc₃₀ nc₃₅ nc 35 nc 40 t₂ ₅ t₂₆t₂ ₃ t₂₄ te ₂₄ t₂ ₅ t₂₆t₂ ₃ t₂₄ te ₂₄ gc-ms(sim) m/z 191 gc-ms(sim) m/z 191 gc-fid gc-fid gc-ms-ms parent–daughter hopanes gc-ms-ms parent–daughter hopanes gc-ms-ms parent–daughter steranes gc-ms-ms parent–daughter steranes ts tm h28 h29 h30 h31 (s +r) h32 (s +r) h33 (s +r) h34 (s +r) h35 (s +r) d 30 m 29 m 30 ts tm h28 h29 h30 h31 (s +r) h32 (s +r) h33 (s +r) h34 (s +r) h35 (s +r) d 30 m 29 m 30 s29 s30 s29 αα αs αα αr αβ βs αβ β r s30 αα αr αα βd27 d27 d27 d27 d27 d27 d27 d27 sample #17895, 37.88 m sample #18086, 216.55 m αα αs αα αr αβ βs αβ β r αα αr αα β αβ βs https://doi.org/10.34194/geusb.v55.8320 http://www.geusbulletin.org/ bojesen-koefoed et al. 2023: geus bulletin 55. 8320. https://doi.org/10.34194/geusb.v55.8320 16 of 24 geusbulletin.org fig. 10 biological marker data from the brorson halvø-1 borehole. characteristic fingerprints of samples representing the upper (sample #18623, 64.23 m, lower volgian) and lower (sample #18889, 224.69 m, kimmeridgian) parts of the drilled succession. gc-ms-ms parent–daughter traces for steranes and hopanes represent the sum of five and nine parent–daughter transitions, respectively. time si gn al time si gn al time si gn al time si gn al time si gn al time si gn al time si gn al time si gn al pris ta ne pris ta ne ph yt an e ph yt an e nc₁₅ nc₁₅ nc₂₀ nc₂₀ nc₂₅ nc₂₅ nc₃₀ nc₃₀ nc₃₅ nc₃₅ nc₄₀ t₂ ₅ t₂₆t₂ ₃ t₂₄ te ₂₄ t₂ ₅ t₂₆t₂ ₃ t₂₄ te ₂₄ gc-ms(sim) m/z 191 gc-ms(sim) m/z 191 gc-fid gc-fid gc-ms-ms parent–daughter hopanes gc-ms-ms parent–daughter hopanes gc-ms-ms parent–daughter steranes gc-ms-ms parent–daughter steranes ts tm h28 h29 h30 h31 (s +r) h32 (s +r) h33 (s +r) h34 (s +r) h35 (s +r) d 30 m 29 m 30 ts tm h28 h29 h30 h31 (s +r) h32 (s +r) h33 (s +r) h34 (s +r) h35 (s +r) d 30 m 29 m 30 s29 s30 s29 s30 d27 d27 d27 d27 d27 d27 d27 d27 sample #18623, 64.23 m sample #18889, 224.69 m nc₄₀ αα αs αα αrαβ βsαβ β r αα αr αα β αα αs αα αr αβ βsαβ β r αα αr αα β αβ βs https://doi.org/10.34194/geusb.v55.8320 http://www.geusbulletin.org/ bojesen-koefoed et al. 2023: geus bulletin 55. 8320. https://doi.org/10.34194/geusb.v55.8320 17 of 24 geusbulletin.org littke et al. 1991; petsch et al. 2000; tang et al. 2018; pan et al. 2022). unravelling the causes for this difference is, however, beyond the scope of the present paper. 5.4 kinextm modelling variations in petroleum generation potential in the shale section of the rødryggen-1 borehole were table 4 key biological marker data from the rødryggen-1 borehole sample # stratigraphy depth (m) pr/ph bhn index d30 index ihr hhi s27 (%) s28 (%) s29 (%) s30 (%) s29 (s/(s+r)) s29 ββ/ (ββ+αα) dia/reg steranes δ13c (total) 17097 ryazanian 25.00 2.52 2.89 7.81 0.12 4.89 32.1 28.3 29.0 10.6 0.30 0.29 2.41   17886 ryazanian 28.89 2.15 7.57 9.35 0.11 6.39 35.9 29.2 24.2 10.6 0.30 0.33 2.16 –29.1 17895 upper volgian 37.88 2.55 2.13 9.22 0.11 4.69 37.1 27.9 27.1 7.9 0.29 0.33 1.96 –29.0 17905 middle volgian 47.06 2.85 1.98 9.68 0.12 4.18 35.4 27.9 26.0 10.7 0.30 0.30 1.77 –29.2 17916 middle volgian 56.88 2.24 4.68 9.73 0.11 6.65 30.6 33.1 26.8 9.4 0.30 0.30 1.61 –29.2 17924 middle volgian 64.47 3.34 7.79 6.01 0.11 4.24 35.1 29.7 26.8 8.3 0.30 0.30 1.72   17935 middle volgian 74.92 3.84 29.10 7.89 0.10 5.12 35.2 31.1 25.7 8.0 0.33 0.32 1.78 –29.9 17944 middle volgian 83.57 2.15 19.42 8.55 0.10 8.18 32.9 31.9 28.5 6.6 0.33 0.32 1.61 –29.7 17953 middle volgian 92.23 2.04 3.12 8.26 0.10 7.16 31.9 31.6 28.6 7.8 0.33 0.29 1.28 –29.9 17962 lower volgian 100.73 2.21 1.40 6.23 0.10 6.42 31.5 31.7 29.0 7.9 0.35 0.30 1.52 –29.4 17973 lower volgian 110.15 2.38 43.26 8.57 0.11 5.54 30.7 31.3 31.5 6.5 0.37 0.33 2.37 –28.4 17983 lower volgian 119.77 3.19 68.03 8.71 0.12 4.23 30.6 29.6 34.7 5.0 0.38 0.36 2.58 –28.1 17992 lower volgian 128.22 3.82 65.55 8.70 0.12 4.20 28.2 28.7 37.1 6.0 0.41 0.23 0.77 –28.2 18002 lower volgian 139.26 3.13 82.51 9.27 0.11 5.53 26.7 17.8 51.0 4.4 0.42 0.24 1.54 –27.3 18012 lower volgian 148.46 3.13 72.91 9.62 0.11 4.51 25.8 21.1 48.7 4.4 0.44 0.23 0.77 –27.2 18023 kimmeridgian 158.77 2.64 70.66 9.81 0.12 4.86 28.7 27.9 37.8 5.6 0.42 0.34 2.18 –27.8 18032 kimmeridgian 167.24 2.70 57.72 7.62 0.11 4.31 29.7 25.3 40.1 4.9 0.44 0.32 2.08 –27.3 18042 kimmeridgian 176.33 2.51 47.73 8.73 0.10 4.63 31.8 27.3 35.4 5.6 0.43 0.32 1.81   18054 kimmeridgian 187.85 4.18 28.12 7.25 0.12 2.85 31.7 27.8 33.3 7.1 0.45 0.33 2.04 –27.3 18064 kimmeridgian 197.24 2.71 46.42 8.33 0.12 3.50 30.4 29.3 34.2 6.1 0.45 0.34 1.83 –27.4 18075 kimmeridgian 207.21 3.07 53.48 8.45 0.11 3.84 30.3 29.2 34.6 5.8 0.49 0.41 2.08 –26.7 18086 kimmeridgian 216.55 3.94 76.54 7.73 0.13 3.41 28.1 27.0 40.8 4.1 0.48 0.38 2.66   18094 kimmeridgian 224.44 2.94 28.86 7.97 0.12 4.20 31.0 28.3 35.2 5.5 0.49 0.35 2.06 –27.0 18109 kimmeridgian 234.66 3.54 78.08 8.37 0.12 4.36 28.2 28.1 37.9 5.8 0.49 0.37 1.59 –26.6 pr/ph: pristane/phytane ratio. bnh index: 100 × (28,30-bisnorhopane/(28,30-bisnorhopane+hopane)). d30 index: 100×*(c30-diahopane/(c30-diahopane+hopane)). ihr: isohopane ratio (nytoft 2011). hhi: homohopane index: 100 × (h35 hopanes)/(sum of h31–35 hopanes), s27 (%), s28 (%), s29 (%), s30 (%). normalised distribution of c27–c30 total steranes, s29 (s/(s+r)). c29 sterane 20s/(20s+20r) isomer ratio, s29 (ββ/ββ+αα). c29 sterane αββ/ (αββ+ααα) isomer ratio, dia/reg. total diasteranes/total regular steranes, δ13c (total). stable carbon isotopic composition, total extract. table 5 key biological marker data from the brorson halvø-1 borehole sample # stratigraphy depth (m) pr/ph h28 index d30 index ihr hhi s27 (%) s28 (%) s29 (%) s30 (%) s29 (s/(s+r)) s29 ββ/ (ββ+αα) dia/reg steranes δ13c (total) 18599 middle volgian 41.28 2.19 0.00 8.73 0.10 12.91 34.0 28.8 25.9 11.4 0.56 0.56 3.30 –30.0 18600 middle volgian 42.09 2.31 0.28 7.63 0.10 12.41 33.2 27.3 26.6 12.9 0.54 0.55 3.41 18611 lower volgian 52.41 2.12 0.51 7.83 0.11 12.12 27.6 29.3 34.6 8.6 0.53 0.55 3.39 20158 lower volgian 52.95 2.10 0.86 7.18 0.11 12.09 28.2 29.0 34.9 7.8 0.55 0.56 3.23 18623 lower volgian 64.23 3.46 22.75 7.87 0.12 9.39 27.7 25.8 39.2 7.3 0.52 0.54 3.81 –28.6 18635 lower volgian 75.28 2.52 19.46 9.87 0.12 10.22 28.0 25.3 39.5 7.2 0.54 0.55 2.73 18648 lower volgian 87.78 2.67 23.49 9.60 0.12 10.01 30.9 24.4 37.2 7.6 0.53 0.56 2.32 –28.8 18659 lower volgian 98.08 2.52 23.91 8.75 0.11 10.26 25.7 23.9 43.6 6.8 0.54 0.60 1.54 –28.0 18672 lower volgian 111.25 2.84 24.48 7.81 0.11 10.73 24.3 19.4 50.0 6.4 0.56 0.61 1.17 –27.8 18695 lower volgian 132.43 2.35 26.22 7.76 0.12 10.12 23.8 18.7 52.2 5.2 0.54 0.60 1.45 –27.4 18711 lower volgian 148.83 2.47 4.03 8.76 0.14 8.75 29.4 25.6 36.8 8.1 0.55 0.61 2.35 18718 lower volgian 155.74 2.75 3.28 9.81 0.14 6.94 29.4 24.3 37.5 8.8 0.54 0.61 3.01 –27.7 19820 kimmeridgian 172.02 2.93 12.36 8.89 0.13 8.19 29.4 24.3 40.5 5.8 0.56 0.63 2.85 18741 kimmeridgian 178.19 2.99 16.32 10.04 0.15 8.49 27.0 20.1 48.2 4.8 0.53 0.62 2.19 –27.8 19821 kimmeridgian 184.35 2.87 10.19 10.86 0.13 7.75 27.9 23.3 42.7 6.1 0.54 0.65 3.13 18865 kimmeridgian 201.23 2.79 9.91 14.21 0.17 8.58 28.3 21.6 44.1 6.0 0.55 0.62 3.34 –26.7 18882 kimmeridgian 217.89 2.38 4.15 20.58 0.19 10.40 30.8 23.1 39.6 6.4 0.52 0.62 1.54 18889 kimmeridgian 224.69 3.07 3.65 21.11 0.18 7.79 26.8 23.7 43.9 5.5 0.52 0.61 4.24 –26.9 pr/ph. pristane/phytane ratio, bnh index. 100 × (28,30-bisnorhopane/(28,30-bisnorhopane+hopane)), d30 index. 100 × (c30-diahopane/(c30-diahopane+hopane)), ihr. isohopane ratio (nytoft 2011), hhi. homohopane index: 100 × (h35 hopanes)/(sum of h31–35 hopanes), s27 (%), s28 (%), s29 (%), s30 (%). normalised distribution of c27–c30 total steranes, s29 (s/(s+r)). c29 sterane 20s/(20s+20r) isomer ratio, s29 (ββ/ββ+αα). c29 sterane αββ/ (αββ+ααα) isomer ratio, dia/reg. total diasteranes/total regular steranes, δ13c (total). stable carbon isotopic composition, total extract. https://doi.org/10.34194/geusb.v55.8320 http://www.geusbulletin.org/ bojesen-koefoed et al. 2023: geus bulletin 55. 8320. https://doi.org/10.34194/geusb.v55.8320 18 of 24 geusbulletin.org fig. 11 biological marker parameters versus drilled depth below reference level (b.rfl.) for the rødryggen-1 (red symbols) and brorson halvø-1 (blue symbols) boreholes (upper panels) and the rødryggen-1 (red symbols) and brorson halvø-1 (blue symbols) boreholes taking into account stratigraphic information (lower panels; alsen et al. 2023). two hiatuses in the succession penetrated by the brorson halvø-1 borehole have been compensated for by assuming that the thickness of the missing section equals the equivalent section in the rødryggen-1 borehole, which shows no hiatus. hence, only samples of the rødryggen-1 borehole show true drilled depths, while samples below the hiatuses in the brorson halvø-1 section have been shifted to greater depths and may even appear deeper than the total depth (td) of the brorson halvø-1 borehole. stratigraphic ages as defined by the rødryggen-1 borehole. quantified by estimating the generated fluid phases (oil and gas) using kinex™ modelling (zetaware inc.; herein referred to simply as kinex modelling) to calculate the ultimate expulsion potential (uep; mmboe/ km2). the uep shows the ultimate expellable volume of hydrocarbons in million barrels of oil equivalent (mmboe) per km2 if the entire shale section cored by rødryggen-1 passed through the petroleum generation window. calculation of the uep applies the kinetics of organofacies b (marine clay-rich shale, largely 0 0.05 0.1 0.15 0.2 0.25 isohopane ratio 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 0 0.5 1 1.5 2 2.5 gammacerane index 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 0 5 10 15 total s30 (%) 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 –31 –30 –29 –28 –27 –26 δ¹³c (total extract) 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 1 2 3 4 5 pristane/phytane 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 m et re s b. r � . 0 5 10 15 homohopane ratio (%) 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 0 0.05 0.1 0.15 0.2 0.25 isohopane ratio 300 290 280 270 260 250 240 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 0 0.5 1 1.5 2 2.5 gammacerane index 300 290 280 270 260 250 240 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 0 5 10 15 total s30 (%) 300 290 280 270 260 250 240 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 –31 –30 –29 –28 –27 –26 δ¹³c (total extract) 300 290 280 270 260 250 240 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 1 2 3 4 5 pristane/phytane 300 290 280 270 260 250 240 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 m et re s b. r � . 0 5 10 15 homohopane ratio (%) 300 290 280 270 260 250 240 230 220 210 200 190 180 170 160 150 140 130 120 110 100 90 80 70 60 50 40 30 20 10 0 kim merid gian l. v olgian m. v olgian u. volgian ryaza nian https://doi.org/10.34194/geusb.v55.8320 http://www.geusbulletin.org/ bojesen-koefoed et al. 2023: geus bulletin 55. 8320. https://doi.org/10.34194/geusb.v55.8320 19 of 24 geusbulletin.org fig. 12 stable carbon isotope ratio (δ13c) of total extracts versus pristane/phytane ratio. note gradual isotopic depletion occurs parallel to decreasing average pristane/phytane ratio up-section. plot modified from chung et al. (1992). fig. 13 comparison of outcrop and borehole data at the rødryggen-1 drill site. the outcrop profile was sampled using the sharp boundary to the albrechts bugt member as datum. orange curve: outcrop data. black curve: drill-core samples. –32 –31 –30 –29 –28 –27 –26 –25 –24 –23 –22 –21 δ¹³c (total extract) 0 1 2 3 4 5 6 pr is ta ne /p hy ta ne ra ti o rødryggen-1, upper volgian lower ryazanian rødryggen-1, middle volgian rødryggen-1, lower volgian rødryggen-1, kimmeridgian brorson halvø-1, middle volgian brorson halvø-1, lower volgian brorson halvø-1, kimmeridgian marine shale and paleozoic carbonate deltaic mesozoic carbonate up section 0 100 200 300 400 500 hydrogen index 60 50 40 30 20 0 4 8 12 16 ts (%) 60 50 40 30 20 0 1 2 3 4 5 6 toc (%) 60 50 40 30 20 d ep th /p os it io n (m ) r el at iv e to c om m on d at um drill-core samples outcrop samples contact to albrechts bugt member https://doi.org/10.34194/geusb.v55.8320 http://www.geusbulletin.org/ bojesen-koefoed et al. 2023: geus bulletin 55. 8320. https://doi.org/10.34194/geusb.v55.8320 20 of 24 geusbulletin.org fig. 14 ultimate expulsion potential (uep) of the cored shale section on rødryggen-1 borehole. the uep profile through the shales shows varying source-rock quality and generation potential in terms of generation products. the shales are relatively gas-prone, but the middle volgian to ryazanian shales are mostly oil-prone. l.: lower. m.: middle. u.: upper. fig. 15 cumulative expelled c6+ hydrocarbons (oil, hc6+) and c1–5 hydrocarbons (gas, hc1–6) of the rødryggen-borehole. oil generation terminates at about 1.0–1.1% ro. equal to type ii kerogen; pepper & corvi (1995a)) and the pepper & corvi (1995b) expulsion model where the retention of hydrocarbons is only adsorption-controlled (hydrocarbons will expel when the adsorption threshold is exceeded). the shale section was split into 258 subsections based on cuttings sampling density, resulting in an average subsection thickness of 0.91 m. due to thermal immaturity of the shales (vitrinite reflectance <0.61%  ro), the measured toc and hi values for each cuttings/subsection were used. the total uep of the c. 235 m thick shale section is c. 44 mmboe/km2, but the source-rock quality varies significantly through the shales (fig. 14). the kimmeridgian and lower volgian sections show the poorest source-rock quality and are relatively gas-prone. the source-rock quality improves significantly in the middle volgian to ryazanian age mudstones that show increased capacity to generate liquid hydrocarbons (fig. 14). however, overall the mudstone section in rødryggen-1 is relatively gas-prone and does not possess the outstanding source-rock quality commonly observed in the upper volgian and ryazanian shales of the kimmeridge clay formation and equivalents, such as the spekk, draupne, mandal and farsund formations in the north sea and atlantic margin (e.g. von der dick et  al. 1989; miller 1990; chakhmakhchev et  al. 1994; klemme 1994; telnæs et  al. 1994; fowler & mcalpine 1995; isaksen & ledje 2001; ineson et al. 2003; justwan & dahl 2005; justwan et al. 2005, 2006a,b; petersen et al. 2010). the uep can be divided into an ultimate expulsion oil (ueo) and an ultimate expulsion gas (ueg), which are c. 15.6 mmboe/km2 and c. 28.6 mmboe/km2, respectively. this is in line with hi values in the most oil-prone middle volgian, upper volgian and ryazanian shales only reaching maximum values of 431 mg hc/g toc, 405 mg hc/g toc and 445 mg hc/g toc. these moderate maximum hi values may reflect the narrow nature of the rift (approx. 30 km), which allows a background contribution of terrigenous organic matter with low potential for petroleum generation, irrespective of basin stagnation and oxygen deficiency. the relatively gas-prone character of the shales is also illustrated by the modelled cumulative expelled volumes where oil (hc6+) generation and expulsion end at about 1.0–1.1%  ro while significant gas generation continues at higher maturities (fig. 15). 6. conclusions the rødryggen-1 and brorson halvø-1 fully cored boreholes were drilled in wollaston forland (northeast greenland, approx. 74°n). they have tested the development in sedimentary environments and petroleum-source potential of upper jurassic – lower cretaceous organic-rich mudstones at two different positions within an evolving halfgraben, situated at the margin of the main north atlantic rift system, and thus partially detached from it. the overall halfgraben is, however, segmented in several subblocks, which evolved quasi-independently through the upper jurassic – lower cretaceous. the two boreholes were thus drilled at different locations within the permpas–hühnerbjerg block(s), bounded l.-u. ryazanian u. volgian m. volgian l. volgian kimmeridgian total uep = 44.2615 mmboe/km² thickness = 234.46 m 0 0.1 0.2 0.3 0.4 0.5 uep (mmboe/km²) 0 25 50 75 100 125 150 175 200 225 th ic kn es s (m ) cumulative hc6+ expelled (mass) cumulative hc1–6 expelled (mass) vitrinite re�ectance easy-ro (%) c um ul at iv e h c e xp el le d (m as s) 1 0.8 0.6 0.4 0.2 0 0 0.5 1 1.5 2 2.5 3 https://doi.org/10.34194/geusb.v55.8320 http://www.geusbulletin.org/ bojesen-koefoed et al. 2023: geus bulletin 55. 8320. https://doi.org/10.34194/geusb.v55.8320 21 of 24 geusbulletin.org fig. 16 summary of the combined stratigraphic coverage of the blokelv-1 (jameson land), rødryggen-1 and brorson halvø-1 boreholes (wollaston forland). for details on the blokelv-1 borehole, see ineson & bojesen-koefoed (2018). combined information from the three boreholes indicates that deposition of organic-rich mudstones with petroleum-generation potential prevailed for a period of approximately 20 million years in north-east greenland during the late jurassic – early cretaceous. u.: upper. m.: middle. l.: lower. by the dombjerg fault to the west and the hühnerbjerg fault to the east, which probably were the main controlling faults in the studied block during the late jurassic. the rødryggen-1 borehole was drilled to a depth of approx. 236 m, in the central part of the basin and includes an uninterrupted mudstone succession ranging in age from the upper kimmeridgian to the upper valanginian. the brorson halvø-1 borehole was drilled to a depth of approximately 226 m close to the uplifted eastern crest of the same block and includes a succession ranging in age from the upper kimmeridgian to the lower barremian, however with two major hiatuses in the middle volgian and upper volgian – ryazanian sections. based on the combined indications of several independent parameters such as tmax, vitrinite reflectance and sterane isomerization ratios, the rødryggen-1 succession is thermally immature but approaching oil-window maturity, whereas the brorson halvø-1 succession is oil-window mature. the causes for the differences in thermal maturity are not clear but may be linked to the presence of abundant palaeogene intrusions in the vicinity of the brorson halvø-1 drill site. such intrusions are not known near the rødryggen-1 drill site. both boreholes show consistent increases in thermal maturity with depth, which is surprising considering the limited depth of the boreholes. the trends may be explained by rapid uplift, which quenched non-linear maturity trends, characteristic of the oil window. a similar trend was reported for the blokelv-1 borehole in jameson land. bearing in mind the hiatus and the differences in thermal maturity between the two boreholes, the organic geochemical characteristics of the two successions are remarkably similar, despite the differences in setting within the graben system. others have demonstrated notable differences in the detailed sedimentology of the two successions, but the impact of sedimentological processes on the organic matter content of the deposits seems to have been trivial in this case. this suggests that the organic matter content of the deposits was governed by higher order processes such as primary productivity and the preservation potential at the sediment-water interface, and that these factors remained rather constant irrespective of the sedimentary processes operating near the seabed. the kimmeridgian through ryazanian succession in general records a transgressive development with gradual decrease in terrigenous organic matter input, increasing bottom-water stagnation oxygen deficiency and marine organic-matter input. superimposed on the c re ta c eo u s (p ar s) lo w er (p ar s) u pp er ju r a ss ic bl ok el v1 rø dr yg ge n1 c. 2 0 m ill io n ye ar s br or so n h al vø -1 o xf o rd ia n l. l. l. l. l. l. m . u . u . l. u . u . u . u . sy st em se ri es th et ys (s ta nd ar d) b or ea l su bs ta ge m . u . ki m m er id g ia n va la n g in ia n be rr ia si a n ti th o n ia n ry a z a n ia n v o lg ia n ba rr em ia n h a u te ri v ia n 130 125 age (ma) 140 150 160 135 145 155 stage https://doi.org/10.34194/geusb.v55.8320 http://www.geusbulletin.org/ bojesen-koefoed et al. 2023: geus bulletin 55. 8320. https://doi.org/10.34194/geusb.v55.8320 22 of 24 geusbulletin.org overall trend are several minor trends related to the detailed development of the rift, discussed in detail by hovikoski et al. (2023a,b). the petroleum-generation potential of the immature/early mature kimmeridgian through ryazanian succession of the rødryggen-1 borehole may be described as oil-prone or gas–oil-prone, growing increasingly oil-prone upwards. however, kinex modelling demonstrates that the succession is relatively gasprone throughout due to the relatively high input of terrigenous organic matter due to the rather proximal setting of the basin compared to the outboard basins of the main rift. a comparison of parallel sets of outcrop and drillcore samples from the rødryggen outcrop and rødryggen-1 borehole, respectively, shows that the hi of outcrop samples may be reduced to one third due to weathering. although the underlying reason for this effect is not documented in full detail, it can be reasonably linked to weathering of abundant, finely disseminated pyrite present in the rocks, the products of which will attack the kerogen and deteriorate its petroleum-generation potential. in summary, the combined evidence from the drilling of the blokelv-1 borehole in jameson land and the rødryggen-1 and brorson halvø-1 boreholes in wollaston forland (fig. 16) indicates that continuous deposition of mudstones with high petroleum-generation potential prevailed in east and north-east greenland over a timespan of approximately 20 million years in the upper jurassic – lower cretaceous. acknowledgments the drilling teams, including john boserup, anders clausen, peter turner, lars (lasse) thomsson, andreas hjort frandsen, anders pilgaard and annette ryge, are thanked for their efforts and for keeping up the good spirits through the very challenging 2009 field season. the pilots finn rusanes and göran lindmark are thanked for diligent sling-work under very difficult conditions. ditte kiel-dühring and carsten guvad are thanked for excellent laboratory work, and hans peter nytoft is thanked for producing biological marker data. jette halskov and annabeth andersen assisted with the draft work. helpful and constructive reviews by drs erdem idiz and iain c. scotchman are gratefully acknowledged. funding statement the present work was funded by geus and a consortium of different companies of the international petroleum industry, who are all partners in a long-lasting collaborative effort, run by geus, concerning the geology of north-east greenland. competing interests the authors declare that they have no competing interests author contributions jabk: writing – original draft, supervision, project administration, investigation, funding acquisition. pa: investigation, funding acquisition. mb: investigation. jh: investigation, writing – original draft. pj: investigation. hnh: investigation. hip: investigation. sp: investigation. hv: investigation. references alsen, p. & mutterlose, j. 2009: the early cretaceous of north-east greenland: a crossroads of belemnite migration. palaeogeography, palaeoclimatology, palaeoecology 280, 168–182. https://doi.org/10.1016/j. palaeo.2009.06.011 alsen, p., piasecki, s., nøhr-hansen, h., pauly, s., sheldon, e. & hovikoski, j. 2023: stratigraphy of the upper jurassic to lowermost cretaceous in the rødryggen-1 and brorson halvø-1 boreholes, wollaston forland, north-east greenland. geus bulletin 55, 8342 (this volume). https:// doi.org/10.34194/geusb.v55.8342 bjerager, m., alsen, p., bojesen-koefoed, j., fyhn, m.b.w., hovikoski, j., ineson, j., nøhr-hansen, h., nielsen, l.h., piasecki, s. & vosgerau, h. 2020: cretaceous lithostratigraphy of north-east greenland. bulletin of the geological society of denmark 68, 37–93. birkelund, t. & perch-nielsen, k. 1976: late palaeozoic – mesozoic evolution of central east greenland. in: escher, a. & watt, w.s. 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(eds): the jurassic of denmark and greenland. geological survey of denmark and greenland bulletin 1, 659– 722. https://doi. org/10.34194/geusb.v1.4644 surlyk, f. et al. 2021: jurassic stratigraphy of east greenland. geus bulletin 46. 6521. https://doi.org/10.34194/geusb.v46.6521 surlyk, f. & clemmensen, l.b. 1983: rift propagation and eustacy as controlling factors during jurassic inshore and shelf sedimentation in northern east greenland. sedimentary geology 34, 119–143. https:// doi.org/10.1016/0037-0738(83)90083-0 surlyk, f. & korstgård, j. 2013: crestal unconformities on an exposed jurassic tilted fault block, wollaston forland, east greenland as an analogue for buried hydrocarbon traps. journal of marine and petroleum geology 44, 82−95. https://doi.org/10.1016/j.marpetgeo.2013.03.009 tang, x., zhang, j., liu,y., yang, c., chen, q., dang, w. & zhao, p. 2018: geochemistry of organic matter and elements of black shale during weathering in northern guizhou, southwestern china: their mobilization and interconnection. geochemistry 78(1), 140–151 http://dx. doi.org/10.1016/j.chemer.2017.08.002 taylor, g.h., teichmüller, m., davis, a., diessel, c.f.k., littke, r. & robert, p. 1998: organic petrology, 704 pp. gebruder borntraeger. telnæs, n., isaksen, g.h. & douglas, a.g. 1994: a geochemical investigation of samples from the volgian bazhenov formation, western siberia, russia. organic geochemistry 21(5), 545–558. https://doi. org/10.1016/0146-6380(94)90105-8  vischer, a. 1943: die postdevonische tektonik von ostgrönland zwischen 74° und 75° n. br. kuhn ø, wollaston forland, clavering ø und angrenzende gebiete. meddelelser om grønland 133(1), 195. von der dick, h., meloche, j.d., dwyer, j. & gunther, p. 1989: source-rock geochemistry and hydrocarbon generation in the jeanne d’arc basin, grand banks, offshore eastern canada. journal of petroleum geology 12, 51–68. https://doi.org/10.1111/j.1747-5457.1989.tb00220.x  https://doi.org/10.34194/geusb.v55.8320 http://www.geusbulletin.org/ https://doi.org/10.1144/0060903 https://doi.org/10.34194/bullggu.v128.6670 https://doi.org/10.34194/bullggu.v128.6670 https://doi.org/10.1016/0031-0182(90)90205-l https://doi.org/10.1016/0031-0182(90)90205-l https://doi.org/10.34194/geusb.v1.4644 https://doi.org/10.34194/geusb.v1.4644 https://doi.org/10.34194/geusb.v46.6521 https://doi.org/10.1016/0037-0738(83)90083-0 https://doi.org/10.1016/0037-0738(83)90083-0 https://doi.org/10.1016/j.marpetgeo.2013.03.009 http://dx.doi.org/10.1016/j.chemer.2017.08.002 http://dx.doi.org/10.1016/j.chemer.2017.08.002 https://doi.org/10.1016/0146-6380(94)90105-8 https://doi.org/10.1016/0146-6380(94)90105-8 https://doi.org/10.1111/j.1747-5457.1989.tb00220.x organic geochemistry of an upper jurassic – lower cretaceous mudstone succession in a narrow graben setting, wollaston forland basin, north-east greenland 1. introduction 2. geological setting and stratigraphy of the drilled succession 3. samples and methods 4. results 4.1 thermal maturity 4.2 petroleum potential and organic facies variations 4.3 outcrop versus borehole samples: weathering effects 5. discussion 5.1 thermal maturity 5.2 petroleum potential and organic facies variations 5.3 outcrop versus borehole samples: weathering effects 5.4 kinextm modelling 6. conclusions acknowledgments funding statement competing interests author contributions references figures fig. 1 location maps. overview map (left). positions of the rødryggen-1 and brorson halvø-1 boreholes on wollaston forland are shown on a simplified geological map (right). k.f.: kuhn fault; p.f.: permpas fault; h.f.: hühnerbjerg fault; d.f.: dombjerg fault. the dombjerg fault was the main fault to control the position of the coastline during the late jurassic. the permpas–hühnerbjerg block(s) was bounded by the kuppel and fig. 2 conceptual, approximately sw–ne-oriented cross-section of clavering ø – wollaston forland, showing multiple westward-tilted fault blocks. modified from birkelund & perch-nielsen (1976) and vischer (1943). fig. 3 stratigraphic models of the wollaston forland – kuhn ø area. left: the original stratigraphy by surlyk (2003). right: the revised stratigraphy after drilling of the rødryggen-1 and brorson halvø-1 boreholes and incorporating changes introduced by surlyk et al. (2021). the revised succession is much more complete and richer in mudstone than suggested by the original. east greenl. tect. strat. seq.: east greenland tectono-stratigraphic sequence. u: upper. m: middle. l: lower. abbreviations for geology are as follows: str: stratumbjerg formation. r: rødryggen member. f: falske bugt member. pa: palnatokes bjerg formation (young sund member). al: albrechts bugt member. n: niesen member. li: lindemans bugt formation. li (s): lindemans bugt formation (storsletten member). ri: rigi member. l: laugeites ravine member. b: bernbjerg formation. ug: ugpik ravine member. j: jakobsstigen formation. pay: payer dal formation. pe: pelion formation. mu: muslingebjerg formation. ba: bastians dal formation. black dashes indicate uncertain contact. black vertical lines indicate boreholes, r-1: rødryggen-1; bh-1: brorson halvø-1. fig. 4 thermal maturity parameters versus drilled depth below reference level (b. rfl.) for the rødryggen-1 (red symbols/lines) and brorson halvø-1 (blue symbols/lines) boreholes (upper panels) and the rødryggen-1 (red symbols/lines) and brorson halvø-1 (blue symbols/lines) boreholes taking into account stratigraphic information (lower panels; alsen et al. 2023, this volume). two hiatuses in the succession penetrated by the brorson halvø-1 borehole have been compensated for by assuming that the thickness of the missing section equals the equivalent section in the rødryggen-1 borehole, which shows no hiatus. hence, only samples of the rødryggen-1 borehole show true drilled depths, while samples below the hiatus in the brorson halvø-1 section have been shifted to greater depths and may even appear deeper than the total depth (td) of the brorson halvø-1 borehole. fig. 5 organic geochemical screening parameters versus drilled depth below reference level (b. rfl.) for the rødryggen-1 (red symbols/lines) and brorson halvø-1 (blue symbols/lines) boreholes (upper panels) and the rødryggen-1 (red symbols/lines) and brorson halvø-1 (blue symbols/lines) boreholes taking into account stratigraphic information (lower panels; alsen et al. 2023, this volume). two hiatuses in the succession penetrated by the brorson halvø-1 borehole have been compensated for by assuming that the thickness of the missing section equals the equivalent section in the rødryggen-1 borehole, which shows no hiatus. hence, only samples of the rødryggen-1 borehole show true drilled depths, while samples below the hiatuses in the brorson halvø-1 section have been shifted to greater depths and may even appear deeper than the total depth (td) of the brorson halvø-1 borehole. stratigraphic ages as defined by the rødryggen-1 borehole. fig. 6 standard plots of organic geochemical screening data from the rødryggen-1 borehole. fig. 7 standard plots of organic geochemical screening data from the brorson halvø-1 borehole. fig. 8 assessment of hilive (i.e. hi of the actively petroleum-generating part of the kerogen) and dead carbon (i.e. the carbon fraction that is inert with respect to petroleum generation) per stratigraphic unit of the rødryggen-1 borehole, following the procedure of dahl et al. (2004). data are listed in table 3. s2: pyrolytic hydrocarbons (mg/g) from rock-eval pyrolysis. toc: total organic carbon. fig. 9 biological marker data from the rødryggen-1 borehole. characteristic fingerprints of samples representing the upper (sample #17895, 37.88 m, upper volgian) and lower (sample #18086, 216.55 m, kimmeridgian) parts of the drilled succession. gc-ms-ms parent–daughter traces for steranes and hopanes represent the sum of five and nine parent–daughter transitions, respectively. fig. 10 biological marker data from the brorson halvø-1 borehole. characteristic fingerprints of samples representing the upper (sample #18623, 64.23 m, lower volgian) and lower (sample #18889, 224.69 m, kimmeridgian) parts of the drilled succession. gc-ms-ms parent–daughter traces for steranes and hopanes represent the sum of five and nine parent–daughter transitions, respectively. fig. 11 biological marker parameters versus drilled depth below reference level (b.rfl.) for the rødryggen-1 (red symbols) and brorson halvø-1 (blue symbols) boreholes (upper panels) and the rødryggen-1 (red symbols) and brorson halvø-1 (blue symbols) boreholes taking into account stratigraphic information (lower panels; alsen et al. 2023). two hiatuses in the succession penetrated by the brorson halvø-1 borehole have been compensated fig. 12 stable carbon isotope ratio (δ13c) of total extracts versus pristane/phytane ratio. note gradual isotopic depletion occurs parallel to decreasing average pristane/phytane ratio up-section. plot modified from chung et al. (1992). fig. 13 comparison of outcrop and borehole data at the rødryggen-1 drill site. the outcrop profile was sampled using the sharp boundary to the albrechts bugt member as datum. orange curve: outcrop data. black curve: drill-core samples. fig. 14 ultimate expulsion potential (uep) of the cored shale section on rødryggen-1 borehole. the uep profile through the shales shows varying source-rock quality and generation potential in terms of generation products. the shales are relatively gas-prone, but the middle volgian to ryazanian shales are mostly oil-prone. l.: lower. m.: middle. u.: upper. fig. 15 cumulative expelled c6+ hydrocarbons (oil, hc6+) and c1–5 hydrocarbons (gas, hc1–6) of the rødryggen-borehole. oil generation terminates at about 1.0–1.1% ro. fig. 16 summary of the combined stratigraphic coverage of the blokelv-1 (jameson land), rødryggen-1 and brorson halvø-1 boreholes (wollaston forland). for details on the blokelv-1 borehole, see ineson & bojesen-koefoed (2018). combined information from the three boreholes indicates that deposition of organic-rich mudstones with petroleum tables table 1 summary of organic geochemical data per stratigraphic interval of the rødryggen-1 borehole table 2 summary of organic geochemical data per stratigraphic interval of the brorson halvø-1 boreho table 3 average contents of inert carbon and hydrogen index values of the reactive kerogen fraction per stratigraphic interval of the rødryggen-1 borehole. calculated using the method of dahl et al. (2004) table 4 key biological marker data from the rødryggen-1 borehole table 5 key biological marker data from the brorson halvø-1 borehole thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 1 of 139 monograph petrology of the skaergaard layered series peter thy,1* christian tegner2 and charles e. lesher1,2 1department of earth and planetary sciences, university of california, davis, usa; 2department of geoscience, aarhus university, aarhus, denmark abstract the skaergaard intrusion is a layered, ferrobasaltic intrusion emplaced during the early eocene into the rifting volcanic margin of east greenland. the magma chamber crystallised in response to cooling from the roof and margins upwards and inward, forming upper, marginal and bottom series, the latter referred to as the layered series. the phase layering in the bottom series suggests an evolved, olivine-normative tholeiitic melt saturated in plagioclase and olivine, followed by augite, and then simultaneously by ilmenite and magnetite forming primocrysts. pigeonite appears in the lower parts and continues until the centre of the series. apatite appears in the upper part concurrently wit h liquid immiscibility. cryptic variations of the individual primocrysts record a systematic upward increase in iron and decrease in magnesium for the mafic minerals and a systematic increase in sodium and decrease in calcium for plagioclase. the appearance of pigeonite is caused by reactions and crystallisation in the trapped melt and by subsolidus adjustments without this phase reaching liquidus saturation. the high mode of olivine at the base of the upper part with the appearance of apatite is interpreted to mark the onset of liquid immiscibility. this may have led to the separation of conjugate melts with granophyre migrating upward and the basic component largely staying stationary or sinking. petrologic and geochemical observations indicate differentiation in the lower part of the intrusion, principally controlled by crystal fractionation with the efficiency of fractionation controlled by the evolution and escape of liquid from the solidifying mush. during the final stages of solidification, the onset of liquid immiscibility and termination of melt convection impeded differentiation. modelling by perfect rayleigh fractionation shows that major and included trace elements conform reasonably to observations, while excluded elements deviate from model predictions. this decoupling is caused by the mobility of a granophyre component formed in the trapped melt and in the main residual magma chamber. consequently, the sampled gabbros may not be representative of the final solid-melt mush. by restoring the gabbros to their original mush compositions, it is possible to constrain granophyre migration pathways. we suggest that the granophyre formed in the trapped melt in the lower part of the intrusion mostly migrated laterally through pressure release pathways to form lenses and pockets with only limited upward migration into the main magma reservoir. near the end stage of differentiation, the residual magma exsolved and formed complex mixtures of ferrobasaltic and granophyric melts. estimates predict that a substantial amount of the granophyric melt penetrated as sills into the downward crystallising, upper part of the body as well as into the host rocks. the redistribution of granophyric melts within the solidifying crystal mush complicates predictions of trapped-melt content and mass-balance calculations but helps to explain apparent decoupling of included and excluded trace elements, especially towards the end stages of evolution. final crystallisation was controlled mostly by in situ crystallisation leaving complex mixtures of ferrodiorite and granophyre components. *correspondence: pthy@ucdavis.edu received: 29 mar 2021 revised: 01 oct 2022 accepted: 22 mar 2023 published: 22 dec 2023 keywords: cumulate, gabbro, greenland, layered intrusion, phase layering abbreviations and symbols 1σ or sd: standard deviation at 68% level an: anorthite as mole per cent (mol%) of plagioclase aug: augite bse: backscattered electron cpx: clinopyroxene d: partition coefficient d*: bulk partition coefficient di: diopside or normative diopside dlc: danish lithosphere centre ds/b: partition coefficient between immiscible silicic (s) and basic (b) conjugate melts en: enstatite or normative enstatite f: fraction of melt remaining feo*, feototal: iron calculated as total oxide feo fmq: fayalite-magnetite-quarts fo: forsterite as mole per cent (mol%) of olivine fo2: oxygen fugacity fs: ferrosilite ftm: fraction of trapped melt germ: geochemical earth reference model geus: geological survey of denmark and greenland gps: global positioning system h: stratigraphic height from the base of the skaergaard profile (m) hz: hidden zone icp-ms: inductively coupled plasma mass spectrometry il: ilmenite iugs: international union of geological sciences iw: iron-wüstite kd: exchange coefficient loi: loss on ignition ls: layered series lz: lower zone lza: lower (sub)zone a lza*: lower (sub)zone a* of the mbs corresponding to lza lzb: lower (sub)zone b lzc: lower (sub)zone c m-sites: structural octahedral sites in pyroxene mbs: marginal border series mg#: magnesium number; mg/(mg+fe*) atomic ratio or % with iron (fe*) calculated as total atomic fe morb: mid-ocean ridge basalt mt: magnetite mw: magnetite-wüstite mz: middle zone n: number of observations, normalised nbo: non-bridging oxygens nno: nickel-nickel oxide ol: olivine or normative olivine opx: orthopyroxene or normative orthopyroxene pa·s: pascal second pfu: per formula unit https://doi.org/10.34194/geusb.v56.8327 https://orcid.org/0000-0002-9267-5798 https://orcid.org/0000-0003-1407-7298 https://orcid.org/0000-0003-4033-4809 mailto:pthy@ucdavis.edu thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 2 of 139 geusbulletin.org pge: platinum group elements pi: pigeonite pl: plagioclase or normative plagioclase q: quartz or normative quartz σr2: sum of the square of coefficients of multiple correlations ree: rare-earth element(s) sh: sandwich horizon t: temperature (°c) t-site: structural tetrahedral site in pyroxene ubs: upper border series uz: upper zone uz’: ubs equivalent to uz uza: upper (sub)zone a uza’: ubs equivalent to uza uzb: upper (sub)zone b uzb’: ubs equivalent to uzb uzc: upper (sub)zone c uzc’: ubs equivalent to uzc wo: wollastonite or normative wollastonite xrf: x-ray fluorescence δ18o: ratio of stable isotopes 18o and 16o δfmq: oxygen fugacity (fo2) normalised to the fmq buffer %rsd: per cent relative standard deviation geus bulletin (eissn: 2597-2154) is an open access, peer-reviewed journal published by the geological survey of denmark and greenland (geus). this article is distributed under a cc-by 4.0 licence, permitting free redistribution, and reproduction for any purpose, even commercial, provided proper citation of the original work. author(s) retain copyright. edited by: jakob kløve keiding (geus, denmark) reviewed by: rais latypov (university of witwatersrand, south africa) and ilya veksler (gfz german research centre for geosciences, germany) funding: see page 117 competing interests: see page 118 additional files: see page 118 https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ https://creativecommons.org/licenses/by/4.0/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 3 of 139 geusbulletin.org 1 introduction basaltic lavas extruded and quenched on the surface of the earth, such as on the sea floor and continental margins and crust, record a complex history of deep-seated melting and crystallisation. these lavas also record the reactive interactions with the host rocks during their upward migration and transient storage in reservoirs, as well as low-temperature water-rock interaction and hydrothermal alteration. because we can sample and analyse extruded lavas, we may be able to deduce their formation and history using various observational, experimental and theory-based modelling approaches (e.g. langmuir et al. 1992; o’hara & herzberg 2002; brown & lesher 2014). despite not being directly observed, temporary storage in crustal chambers has proven necessary for explaining the observed differentiation pattern in many extruded magmas. an important source for our understanding of the role of such magma chambers and related chamber processes has come from studies of tectonically uplifted sea-floor fragments (or ophiolites; e.g. coleman 1977; pallister & hopson 1981; thy et al. 1989; thy & dilek 2000, 2003; bédard 2015). additional information has come from seismic studies of active centres of sea-floor spreading (sinton & detrick 1992; morgan et al. 1994) and from direct drilling into ocean crust near spreading centres (natland et al. 1991; dick et al. 2000; thy 2003). the chamber view that has emerged for steady-state settings of sea-floor spreading is one in which the chamber extends downwards as a zone filled with crystal mush, centred under the spreading ridge and capped by a smaller magma lens maintained by upward migrating fresh melt and interstitial melt migration aided by compaction of the crystal mush (nisbet & fowler 1978; sinton & detrick 1992; kelemen et al. 1997). flux of fresh, mantle-derived magma upwards in the mush may stabilise as sill-like bodies forming sill complexes (kelemen et al. 1997). the notion of such transitional crystal-mush zones beneath volcanic systems have been heralded as a unifying conceptual model that can explain features of both axial and off-axial volcanic systems (cashman et al. 2017; sparks et al. 2019). it is thought to explain, on the one hand, a slow, steady incremental influx of melt, originating deeper in the crust or mantle and building up large mush reservoirs, and on the other hand, episodic large-volume eruptions in the form of plateau lavas. the observation that large-volume magma chambers are often not detected in the crust leads to the suggestion that magmatic melt is, instead, stored in mush reservoirs, only to be activated just prior to eruption on the surface or accumulated in shallower chambers if direct gateways to the surface are not easily available (edmonds et al. 2019; holness et al. 2019; sparks et al. 2019). shallow melt emplacement and trapping in melt-dominated chambers are important building blocks for crustal accretion (bachmann et al. 2007; edmonds et al. 2019). although systematic studies of ophiolitic ultramafic and gabbro complexes are relatively few (thy & dilek 2000, 2003), those which do exist reveal a range of magma chambers. these include a steady-state, seafloor spreading-related chamber that may represent a crustal-mush zone (pallister & hopson 1981; kelemen et al. 1997), backarc-related chambers (thy et al. 1989) and continental rift chambers (church & riccio 1977; bédard 2015) that record infrequently replenished, melt-dominated chambers and thus magmatic settings controlled by restricted upward flux of melt. an equally important source for understanding basaltic magma chambers has been melt-dominated plutons intruded into continental settings. most prominent of these, among many, have been the stillwater complex of montana (hess 1960; jackson 1961; page 1979; mccallum 1996), the kiglapait intrusion of the anorthositic nain complex in labrador (morse 1969, 2015), the rhum (or rum) complex of the inner hebrides, scotland (brown 1956; tait 1985), the ultramafic jurassic-cretaceous duke island complex of south-eastern alaska (irvine 1974), the precambrian muskox intrusion of the canadian northwest territories (irvine 1980), the synorogenetic fongen-hyllingen intrusion of the central norwegian caledonides (wilson et al. 1981; wilson & larsen 1985; wilson & sørensen 1996), the precambrian bushveld complex in the transvaal of south africa (daly 1928; eales & cawthorn 1996; cawthorn 2015) and finally the tertiary skaergaard intrusion of east greenland (wager & deer 1939; wager & brown 1967; mcbirney 1995, 1996; brooks 2011; holness et al. 2017a). this latter intrusion was emplaced into continental crust during the rift-todrift transition leading to formation of the north atlantic ocean basin and is the focus of the present study (fig. 1). these ultramafic to gabbroic plutons represent wide ranges in terms of initial melt composition, volume, depth of intrusion and structural and tectonic history of emplacement. there are, however, some commonalities that often characterise their solidification and cooling histories. typically, the plutons reveal overall internal systematic differentiation patterns that correlate with the inferred decreasing magma temperature (t), often from the base to the top of a melt-dominated stratigraphy. in many cases, the overall differentiation trend is interrupted and reversed for https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 4 of 139 geusbulletin.org shorter periods due to magma recharge. additionally, they display complex layering defined by modal and size variations of the constituent minerals. such plutons are typically melt-dominated and are referred to as layered intrusions. differentiation is recorded in the mineral assemblages (phase layering) with the minerals and modal proportions constrained by known basaltic phase equilibria. additionally, differentiation is recorded in mineral compositions (cryptic layering) and shows systematic changes in sodium (plagioclase) or iron (olivine and pyroxenes) contents increasing with decreasing temperature and fraction of melt remaining (f) upward or inward in the plutons. the observed differentiation is mainly attributed to fractional crystallisation due to cooling, but may also be related to magma mixing, svalbard skandinavia uk greenland iceland charlie-gibbs fracture zone north sea c al ed on ia n fr on t ca le do ni an f ro nt barents s ea arctic ocean vøring plateau rock all p lateau faroe islands disko gak ke l ridge knipovich ridge mohns r idge reykjanes r idge atlantic ocean 80°n 60°n 70°n 30°w 15°w 0° 50°n skaergaard intrusion onshore basalts o�shore basalts seaward-dipping re�ectors dsdp 38 dsdp 81 odp 407-9 odp 104 odp 914-7, 989-90 odp 918 odp 919 deep-sea drilling site 500 km tfz jmfz sisz fig. 1 north atlantic igneous province as outlined by sea-floor spreading ridges, major fracture zones and continental-bordering areas. shown are onshore and offshore basalts, seismic dipping reflector sequences, the caledonian fronts and deep-sea drilling sites, prefixed odp (ocean drilling programme) or dsdp (deep-sea drilling programme). tfz: tjörnes fracture zone; jmfz: jan mayen fracture zone; sisz: south iceland seismic zone. location of the skaergaard intrusion is marked at 68°10´ n, 31°40´ w. map is modified from brooks (2011). https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 5 of 139 geusbulletin.org assimilation of country rocks, liquid immiscibility and fluid exsolution. the separation of crystals from the bulk melt is widely believed to either be due to (1) gravitational accumulation and compaction, (2) crystallisation on a melt-solid interface (hard ground), (3) crystallisation in a bottom-mush zone (soft ground) or some combination of these processes. regardless, the result is the formation of melt-solid interface mushes made of liquidus primocrysts (often referred to as cumulus phases) and variable proportions of interstitial crystals and crystal rims that grew from interstitial melts. the latter two options are referred to as in situ crystallisation processes, where a solidification boundary melt is completely or partially recycled back into the main magma chamber where differentiation is driven by convection and homogenisation. in contrast to extruded lavas, the rocks of slowly cooled intrusions do not preserve the melt from which they crystallised. moreover, melts in the mush can crystallise, migrate or react with host minerals during growth and cooling. thus, in situ crystallisation, compaction and melt convection or migration and reactions all contribute to the demise of interstitial melt. however, their various contributions are challenging to detect and even more difficult to quantify. it is the purpose of this study to review much of what is understood about such differentiation processes as exemplified by the skaergaard intrusion (fig. 1). we present petrographic, mineralogical and major and trace element data of the main layered series (ls) collected during ground traverses and from a drill core. this suite of samples constitutes a composite profile through the ls (tegner et al. 2009) that closely follows the previously constructed profiles through the central portion of the ls by wager & deer (1939), wager & brown (1967) and mcbirney (1989a). we begin with a brief overview of the skaergaard intrusion, followed by a summary of methods and then a description of the petrography upon which the internal structure or zonal divisions are based. from there we examine mineral and bulk-rock variations that form the observational basis for petrologic modelling. in the following sections, we synthesise the great volume of previous work and use our new data to constrain the liquid line of descent and to evaluate proposed and potentially competing models of differentiation for the skaergaard intrusion. a set of appendices are provided at the end of this manuscript that includes additional historical background (appendix 1), information on sampling and construction of the composite profile (appendix 2) and analytical methods and mode determinations (appendix 3). included as online supplementary files is a collection of data tables, photomicrographs and summaries of modelling results. it is our hope that future researchers working on the skaergaard intrusion and similar plutons will find these compilations useful in their own research and that the archive will be a valuable resource for courses in petrology and geochemistry. to better facilitate the latter, we also include the original sample location maps by l.r. wager and a.r. mcbirney, geological and topographical maps, together with a set of stereoscopic aerial colour photos of the skaergaard intrusion and surroundings as online supplementary files. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 6 of 139 geusbulletin.org 2 the skaergaard intrusion the skaergaard intrusion is a relatively small (c. 300 km3) melt-dominated pluton that has been repeatedly and extensively studied because of the field exposure, due to glacial erosion, of a near complete and uninterrupted sequence of fractionated basaltic gabbros. the common explanation is that the skaergaard magma chamber was emplaced as one pulse or a series of rapid pulses at a high level in the east greenland crust and cooled relatively quickly with little interaction with its host rocks (approximating a closed system). the intrusion was discovered in 1930 in the kangerlussuaq fjord of east greenland by l.r. wager and described in detail by wager & deer (1939). since then, the intrusion has provided exceptional insights into the crystallisation and solidification of basaltic magma chambers (e.g. wager & brown 1967; mcbirney 1995; irvine et al. 1998; holness et al. 2017a). a summary of the history of the exploration of the intrusion can be found in appendix 1; for more details the reader is referred to brooks (2018). the skaergaard intrusion intruded at 68°10’n and 31°40’w during continental rifting of the east greenland continental margin (fig. 1; brooks 2011) at a high crustal level at the contact between precambrian gneisses and overlying eocene plateau basalts, which preserve a well-defined metamorphic contact aureole and a chilled margin (wager & deer 1939; hoover 1989a; kays et al. 1989; bufe et al. 2014). the chamber was initiated by small pulses of magma, followed by a large influx that inflated the chamber to its maximum (holness et al. 2015), and this resulted in the present box-shaped intrusion controlled largely by normal faulting (irvine 1991; nielsen 2004). the intrusion was emplaced around 55.4–56.6 ma (hirschman et al. 1997; tegner et al. 2008; brooks 2011; wotzlaw et al. 2012), coinciding with the formation of the main plateau lavas (larsen et al. 1989; tegner et al. 1998; larsen & tegner 2006) and the transition from rift to drift tectonics of the north atlantic (storey et al. 2007; tegner et al. 2008; brooks 2011). as a result of initial cooling, a marginal gabbro (including a chilled margin) formed against the country rocks. the subsequent crystallisation proceeded upward and inwards from the margins of the chamber (fig. 2a). the marginal border series and upper border series (mbs and ubs, respectively) formed by crystallisation from the walls and roof of the chamber (naslund 1984; hoover 1989a, b; salmonsen & tegner 2013), respectively, and the layered series (ls) formed concurrently sh lza hz lzb lzc mz uza uzb uzc + feti oxides + olivine – olivine + apatite + augite mbs ubs ls plagioclase + olivine a lza lzb lzc mz uza uzb uzc lza lzb lzc mz uza uzb uzc th is s tu dy central column b 0 200 400 600 800 1000 1200 1400 1600 1800 2000 2200 0 20 40 60 0 20 40 60 plagioclase olivine augite 0 0.4 0.8 an mol% fo mol% mg/(mg+fe*) st ra ti gr ap hi c he ig ht (m ) wager & deer (1939) wager & brown (1967) mcbirney (1987) c lzc mz lza lzb uza uzb uzc w ag er & b ro w n (1 96 7) fig. 2 summary of the skaergaard intrusion and its layered series (ls). a: schematic representation of the skaergaard intrusion showing the main divisions of the ls after mcbirney (1995) and nielsen (2004). mbs: marginal border series. ubs: upper border series. hz: hidden zone. lz: lower zone (with subdivisions a to c). mz: middle zone. uz: upper zone (with subdivisions a to c). sh: sandwich horizon. the hz and lza have been grouped together. the presence (+) and absence (–) of the main minerals used to define the zone divisions are indicated following wager & brown (1967). b: comparison of the central profile as measured by wager & brown (1967) and as used in this study. c: the cryptic variation in the central profile of the ls as seen for plagioclase (an mol%), olivine (fo mol%) and augite (mg/(mg + fe*) ratio) on an atomic basis with all iron calculated as total ferrous iron. data from wager & deer (1939), wager & brown (1967) and mcbirney (1989a). all data have been projected onto the stratigraphic column of this study. the solid curves are the best visual fit to the data of wager & brown (1967), extended to the full stratigraphic occurrences irrespective of primocryst status. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 7 of 139 geusbulletin.org upward from the floor of the chamber (mcbirney 1989a). the crystallisation fronts eventually converged at the position of the sandwich horizon (sh). as a result, three differentiated stratigraphic sections formed in different positions relative to gravity (figs 2a, b) that allow detailed assessments of the solidification processes (wager & deer 1939; wager & brown 1967; mcbirney 1995). it has been proposed that candidate liquids for the initial skaergaard magma occur as coeval dykes intruding basement gneiss in the vicinity of the intrusion (nielsen 1978) and are found among the plateau lavas of east greenland (larsen et al. 1989; tegner et al. 1998; andreasen et al. 2004; larsen & tegner 2006; jakobsen et al. 2010: nielsen et al. 2019a; cho et al. 2022). the principal manifestation of the evolution of the parental magma is a systematic mineralogical enrichment in low-temperature components or end members, like albite and fayalite, of the constituent minerals upward and towards the centre of the intrusion (fig.  2c). an equally important manifestation is the systematic appearance and disappearance of minerals (phase layering) with falling temperature, which can be related to crystallisation behaviour in simple basaltic systems (osborn 1959; presnall 1966; irvine 1970a, 1979; ford 1981). this allows the layered and border series to be divided into primocryst zones and subzones, defined by the presence or absence of minerals (figs 2a, b). these zonal divisions provide convenient reference horizons for tracking magmatic evolution within the chamber (wager & brown 1967). the prominent role of the intrusion in shaping modern petrologic theories is not only due to the intrusion’s unique coastal glacial-shaped exposure, but perhaps more importantly due to the fundamental compositional and petrographic information that l.r. wager and his co-workers painstakingly collected through the central parts of the intrusion over nearly half a century of work (fig. 2c). wager & deer (1939) constructed their original stratigraphic profile through the ls of the intrusion using approximately only eight samples to define the cryptic mineral compositional variation (fig. 3b). wager & brown (1967) later produced a more detailed record (fig. 3b). a drill core obtained in 1966 penetrated about 150 m into the hidden zone (hz) and extended the central profile to slightly deeper levels (fig. 3a; maaløe 1976a; nwe 1976; holness et al. 2015; brooks 2018). mcbirney (1989a) defined three laterally positioned parallel profiles (west, central and east) of which the central profile was located close to the original profile constructed by wager & deer (1939). a summary of the cryptic mineral variation defined by these early studies in the central column of the intrusion and re-scaled to the stratigraphic column of this study is illustrated in fig. 1c (wager & deer 1939; wager & brown 1967; mcbirney 1989a). the opportunity to re-sample the central profile of the skaergaard intrusion was offered following the decision by the geological survey of denmark and greenland (geus) to salvage and transport to copenhagen some of the most important exploratory drill cores made in 1993 by platinova in the upper part of the intrusion. in 2000, the danish lithosphere centre (dlc) conducted field operations in the area, which offered an opportunity for the retrieval of the drill cores as well as for conducting additional field sampling (nielsen et al. 2000; tegner et al. 2009). this development followed year-long debates on the liquid line of descent of the intrusion, the importance of liquid immiscibility and causes of mineralisation. the present study is an outcome of these opportunities and discussions. the main petrologic interest in the skaergaard intrusion has traditionally been, and still is, in the crystallisation and solidification modes and the resulting liquid fractionation trend. more recently, interest has shifted towards the precious metal (pge-au; where pge refers to platinum group element) mineralisation, which was discovered around 1990 in the upper part of the intrusion (bird et al. 1991) and has since been studied by several groups (e.g. andersen et al. 1998; nielsen et al. 2015, 2019b; rudashevsky et al. 2023). the pge-au mineralisation occurs within a narrow stratigraphic interval reaching approximately 40 m in the centre of the upper part of the ls. given that the mineralisation itself had no resolvable influence on the crystallisation trends for the major element and the trace elements of concern in this study, we do not discuss this aspect of the intrusion further. instead, we encourage interested readers to consult the papers cited above. this study presents the petrographic and petrological results of the detailed sampling (fig. 3) of the central profile of the ls based on mineral exploration drill cores and surface samples precisely located using the global positioning system (gps; tegner et al. 2009) together with new and improved geological (mcbirney 1989b) and topographic maps and aerial photos (see supplementary files s6–s8). our composite profile includes sampling at an average interval of 16 m for a total of 136 samples (fig. 3d). this spacing provides better resolution than previously possible for the petrography, mineral zone divisions, mineral chemistry, density relations, bulk-gabbro compositions (major, minor and trace elements) and their petrologic implications and bearings on the crystallisation of the parental magma. the field relations and internal rhythmic modal-layering structures of the intrusion have been discussed in detail in the original descriptions (wager & deer 1939; wager & brown 1967) and several subsequent studies (maaløe 1978; mcbirney & noyes 1979; naslund & mcbirney 1996; boudreau & mcbirney 1997; irvine et al. 1998; nielsen 2004) and so need only brief mentioning in the context of this petrographical study. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 8 of 139 geusbulletin.org u tt en ta l p la te au kr ae m er s ø fo rb in de ls es gl et sc he r dr ill c or e 90 -2 2 ba si st op pe n lza lzb lzc mz uzb uza uzc ubs 2500 2000 1500 1000 500 0 st ra ti gr ap hi c he ig ht (m ) composite pro le d eoh 1052.47 m l-0 l-1 l-2 600 700 800 900 1000 500 100 0 200 300 400 500 sample granophyric dyke dyke rock mesocratic gabbro (40–65% plagioclase) leucocratic gabbro (>65% plagioclase) troctolite (<10% pyroxenes) core 90-22 c wager & brown (1967) wager & deer (1939) uz mz lz–mz lz this study mz lzb lzc lza m bs uza uza uzc uzb uzb bs mz mz ubs 2 km uttental plateau kraemer ø forbindelsesgletscher basistoppen basisgletscher skærgårdshalvø core 90-22 31°45ʹw 31°35ʹw 68 °0 8ʹ n 68 °1 3ʹ n b uttental sund forbindelsesgletscher basistoppen basement gneiss lzb lzc lzc lza m bs uza uza basalt uzcuzb uzb mz mz ubs uttental plateau kraemer ø cambridge core 2 km a core 90-22 basisgletscher 31°45ʹw 31°35ʹw 68 °0 8ʹ n 68 °1 3ʹ n skærgårdshalvø fig. 3 location of samples and the drill core used in the present study. a: geological map after mcbirney (1989b). b: locality map for the present study as well as the approximate localities used for the profiles constructed by wager & deer (1939) and wager & brown (1967). the grey dashed line marks the profile of this study. the central profile constructed by mcbirney (1989a; not shown) largely follows the general trend of wager & deer (1939). c: drill core 90-22 with dominating petrography and sample positions shown (watts, griffis & mcouat ltd. 1991). l-0, l-1 and l-2 refer to the triple group of prominent felsic layers and the host of the skaergaard pge-au mineralisation. d: construction of the composite profile after tegner et al. (2009). eoh: end of hole. other abbreviations in fig. 2. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 9 of 139 geusbulletin.org 3 methods this chapter briefly summarises the methods of the study, including sampling and construction of the reference column and the analytical techniques. details can be found in the accompanying appendix 2 (sampling) and appendix 3 (analytical methods) as well as in previous papers that used the results of this study (thy et al. 2009a, b; tegner et al. 2009). the bulk gabbro, major element analyses used in this paper were also partly included, together with some trace elements, in a study of the trapped melt content by tegner et al. (2009). the complete data set, together with other available skaergaard analyses, are included in tegner et al. (2023). subsets of the samples were also used in isotope studies by hagen-peter et al. (2019; sr) and lesher et al (2023; fe). the new composite profile (fig. 3) was first described in tegner et al. (2009). it was constructed using a mineral exploration drill core penetrating the upper zone (uz; figs 3c, d) from the lower part of uzc and into the upper 100 m of the middle zone (mz). this core section was supplemented by surface samples (figs 3a, b), extending the coverage of the ls into the lower zone (lz) and the uzc, including the sh and the adjoining part of the ubs. the ubs is not a part of this study. the 51 representative core samples used in the study (fig. 3c) were sampled in 1993 from platinova core 90-22 and represent 1052 m of mostly uz gabbros (watts, griffis & mcouat ltd. 1991; andersen 1996; tegner 1997; andersen et al. 1998). the full drill core, together with other similar cores, have now become available for research purposes at the natural history museum of denmark (nielsen et al. 2000; hanghøj 2005). a total of 85 surface samples were collected in 2000 from profiles largely paralleling the original sampling of wager & deer (1939). the surface samples were located using gps with the purpose of extending the stratigraphic coverage of the drill core samples into the lower and the upper parts of the intrusion. these two sets of samples were used to construct the combined stratigraphic column with an average sample interval of about 16 m (tegner et al. 2009). the locations of the current samples and drill core are summarised in fig. 3b, together with the locations of the samples used by wager & deer (1939) and wager & brown (1967). the zone boundaries in the stratigraphic column established in this study are based on field sampling and petrographic observations, and largely confirm the zone boundaries mapped by mcbirney (1989b) to within the resolution in the constructed composite profile of fig. 3d. this is particularly the case for the appearances of abundant pyroxene (lzb) and feti oxides (lzc), but marginally so for zone boundaries that rely on petrographic interpretations, like mz and uz. however, significant differences exist in the relative thicknesses of some zones and subzones compared to those defined by previous results by mcbirney (1989a) and wager & brown (1967; fig. 2b). most prominent is a reduced thickness for mz that can, at least in part, be ascribed to the improved positioning of samples and zone boundaries by gps. the thicknesses obtained in this study apply to the established central profile but may not apply to averages for the intrusion as a whole or to marginal positions (wager & deer 1939; nielsen 2004). the differences in the various estimates are probably caused by the bowl-shaped zones that may result in an increased average thickness of the mz compared to the present estimates for the central part of the intrusion (irvine 1991; nielsen 2004). with knowledge of the absolute volume relations of the zones and subzones and their corresponding stratigraphic height (h, in m), fraction of remining melt (f) can be predicted as a function of h. the subzone volumes and their mass equivalents were reported by nielsen (2004). a conversion of h to f for the composite profile was proposed by tegner et al. (2009) using a second order polynomial equation as: f = (1.091 × 10–7) × (h2 – 5.9064 × 10–4) × (h + 0.7678) (1) with the result that f = 0.76 at h = 0 m, corresponding to the base of lza in the composite profile. further, eq. (1) predicts that f = 1 at 368 m below lza, or at the base of the intrusion (hz) and are used as the baseline for forward modelling in this study. tegner et al. (2009) assumed that f = 0 at the top of the uzc, corresponding to a height of 2165 m in the composite profile. we also adopt this constraint in our models, although it is not required that the melt will be exhausted on solidification of the ls. values for f for each sample of the composite profile are shown in appendix 2. samples were ground to a fine powder using either corundum (surface samples) or tungsten-carbide (drill-core samples) shatter boxes. for the tungsten-carbide shatter box, significant contamination was only observed for tungsten. loss on ignition (loi) was determined at 950°c and the ferrous-ferric ratios were determined by titration. major elements and selected trace elements were determined by x-ray fluorescence (xrf) methods using a panalytical pw2400 x-ray spectrometer at aarhus university and fused glass (major elements) and powder (trace elements) pellets. the precision (per cent relative standard deviation, https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 10 of 139 geusbulletin.org %rsd) is typically below 1–2% for major elements, 2–4% for minor elements and 1–9% for trace elements. trace element abundances were determined by inductively coupled plasma mass spectrometry (icp-ms) using an agilent 7500 instrument at university of california, davis. precision was evaluated using multiple analyses of reference standards to be well below 6% for most elements. bulk-rock specific density was determined by weighing the dry and clean samples in air and in water with an estimated precision 1–2% (tegner et al. 2009). the mineralogy and texture of both surface and drill-core samples were characterised by optical microscopy. throughout this paper, orthopyroxene refers to low-ca pyroxenes, regardless of whether they precipitated directly as an orthorhombic pyroxene, as may have been the case in the lowermost parts of the intrusion (wager & brown 1967; nwe 1976) or formed from inversion of monoclinic pigeonite. clinopyroxene (cpx) refers to monoclinic calcic and subcalcic pyroxenes (poldervaart & hess 1951). ferrobustamite refers to a pyroxenoid phase that is believed to have been formed by inversion from ferrowollastonite (brown & vincent 1963) resulting in a mosaic of small grains (naslund 1984). the bustamite phase was previously described as ferrowollastonite (brown & vincent 1963). mineral compositions were determined by electron microprobe analysis using either cameca sx-50 or sx-100 instruments (wavelength dispersive spectrometers, university of california, davis) or a jeol 8600 superprobe instrument (combination of wavelength and energy dispersive spectrometers, aarhus university). beam conditions and beam diameters were similar for both laboratories. the beam diameters were either focused to 1–2 µm or defocused or rasterised to a 20–30 µm broad beam to minimise the effects of exsolution in pyroxenes and oxide minerals. an internal pyroxene standard was analysed at both laboratories concurrently as a control and measure of analytical precision and accuracy. except for a few pyroxene batches, no inter-laboratory corrections were needed. precision was estimated to be below 1% for sio2, mgo and cao; below 8% for al2o3, feo, na2o and k2o and below 20% for tio2 and mno. the analyses of apatite included f (fluorine) and cl, but with oh calculated by balance. modal proportions on a weight basis of the constituent minerals were calculated by weighted, least squares, linear approximations using the mineral and the bulk-rock compositions. phases considered in the calculations are olivine, plagioclase, clinopyroxene, orthopyroxene, ilmenite, magnetite and apatite for the entire ls with orthoclase and quartz included for the uz. only acceptable positive solutions to the least squares approximations were compiled with residual sum of squares typically below 0.1. because the average mineral compositions used for the modal calculations represent core compositions, the obtained modal results may be biased by not fully representing late-stage crystallised mineral. such an effect, however, is considered minimal because of the mostly low content of interstitial minerals and because these often are relatively similar to core primocrysts, except in extreme cases of zoning. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 11 of 139 geusbulletin.org 4 stratigraphy of the layered series (ls) 4.1 petrography and zonal division it is customary to refer to the gabbros of the skaergaard intrusion and other layered intrusions as cumulates based on cumulus (or primocryst) mineral assemblages. a consequence is that their interpreted primocryst assemblages classify such gabbros without including an often-large amount of minerals crystallised from an interstitial melt component. a further complication with the cumulus and cumulate terminology is that traditionally it has involved genetic assumptions of gravitative or flow accumulations of cumulus minerals in a host silicate melt (wager & brown 1967). to avoid such genetic interpretation at the onset of this study, we refer to the gabbros as made up of primocrysts and interstitial material, irrespective of mode of formation. the most obvious challenge to the cumulus terminology for skaergaard is that the lowermost gabbros will be classified as troctolite cumulates based on the presence of primocryst plagioclase and olivine, despite them being gabbros or gabbronorites, containing significant amounts of pyroxenes originating from an interstitial melt. during this study of a large number of gabbros, we found it to be impractical to identify the mineral textures as either primocryst or interstitial. for this reason, and to avoid often-subjective interpretations of textures, we use the calculated mineral modes of the bulk compositions to classify the gabbros in accordance with the recommendations of international union of geological sciences (iugs) subcommission on the systematics of igneous rocks (le maitre 1989). by doing so, we conform to the original practice of wager & deer (1939) of classifying the rocks by bulk-mineral modes. one consequence of this practice is that our classification tends to be biased towards the near-solidus phase assemblage rather than the subliquidus assemblage, which in practice elevates the abundance of clinopyroxene and orthopyroxene that are subject to subsolidus exsolution and coarsening as a function of cooling. like the modal-based classifications, the cipw normative compositions calculated from the bulk rocks are predominantly olivine-hypersthene normative. the only exception is the upper part of uzc, where quartz-normative compositions (4–10% quartz) prevail due to a high content of interstitial granophyre. further, for convenience we refer to all clinopyroxeneand plagioclase-bearing rocks as gabbros, despite the fact that from mz and up they are strictly diorites, given that the an content of plagioclase falls below 50% (le maitre 1989). the examined gabbros from the drill core or the field were intended to represent average gabbros. for this reason, we do not use modifying prefixes, such as leuco and melano. for descriptive purposes, we retain the traditional zonal division based on the predominant primocryst mineralogy of wager & deer (1939), wager & brown (1967) and mcbirney (1989a, b, 1996), but prefer the term primocryst instead of cumulus as advocated by maaløe (1976a). the lz is thus characterised by plagioclase, olivine and clinopyroxene; the mz by plagioclase and clinopyroxene, concurrent with reduction in olivine; and the upper zone (uz) by plagioclase and clinopyroxene, while olivine reappears. subzones are further defined by abundant and textural change of clinopyroxene from an interstitial habit to a more dominating granular habit (lzb), feti oxide minerals composed of an intergrowth of ilmenite and magnetite (lzc), euhedral apatite (uzb) and a mosaic intergrowth of green ferrohedenbergite grains interpreted to have inverted from ferrobustamite (uzc). orthopyroxene is present in all zones and subzones, including the mz and into the lower part of uzb. minerals in trace amounts as well as late-stage magmatic, hydrothermal and metamorphic minerals (e.g. hydrous minerals and sulphides) are not considered in the present discussion (e.g. wager et al. 1957; wager & brown 1967; bird et al. 1986, 1988; manning & bird 1986; rudashevsky et al. 2004, 2023; cabri et al. 2005; nielsen et al. 2015, 2019b). we begin by reviewing the petrography of the ls gabbros, as it bears on the zone and subzone divisions of the intrusion, and then consider the stratigraphic variations. fig. 4 provides an overview of modal compositions illustrating the elevated olivine and apatite contents of uzb and uzc and the feti oxide content for much of lzc. zonal averages are provided in table 1. representative plane-polarised, transmitted, microscope images of thin sections of the investigated gabbro samples are given in supplementary files s4. 4.1.1 lower zone (lz) the lz comprises 876 m of stratigraphy within the composite profile (fig. 3) that consist predominantly of gabbronorite (fig. 4) characterised by plagioclase, olivine, clinopyroxene, orthopyroxene and feti oxides. the lz has been subdivided into three subzones (lza, lzb and lzc; fig. 3). the 1966 cambridge drill core i was located in the stratigraphically lowest part of lzb and penetrated through the lza and 150 m into the hz to a total stratigraphic thickness of 1026 m (maaløe 1976a; holness et al. 2015). wager & deer (1939) referred to the gabbros of the lz as “hypersthene-olivine-gabbros” of which the gabbros in the lower-most part (lza) were identified as being made up of “transitional rocks” in the sense that they transitioned into the border group. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 12 of 139 geusbulletin.org 4.1.1.1 lower (sub)zone a (lza) the lza includes 174 m of olivine gabbronorites (fig. 3) exposed at the surface and c. 150 m into the subsurface as documented in cambridge drill core i (maaløe 1976a; nwe 1976; holness et al. 2015). on average, lza has 56 wt% plagioclase, 9 wt% olivine, 13 wt% clinopyroxene and 18 wt% orthopyroxene (table 1; figs 4, 5) and contains medium-grained (2–6 mm) primocrysts of plagioclase and olivine predominantly forming a granular network in which olivine tends to be rounded to subhedral and plagioclase tabular in a random to orientated fabric. ilmenite, magnetite and apatite occur interstitially and rarely exceed a total of 4 wt% (fig. 6). biotite is rare and often intergrown with apatite or associated with feti oxides. orthopyroxene occurs as coronas between feti oxides and olivine or plagioclase (fig. 6; haselton & nash 1975; holness et al. 2007a, b). plagioclase is strongly twinned and displays complex normal and polysynthetic twinning in the lower part of lza, becoming more homogeneous higher in the stratigraphy (wager & brown 1967; maaløe 1976a; humphreys 2009; namur et al. 2014). plagioclase primocrysts also occur as small (<1 mm) chadacrystic grains in large oikocrystic clinopyroxene. 4.1.1.2 lower (sub)zone b (lzb) the lzb comprises 542 m of gabbronorites (fig. 3) containing on average 51 wt% plagioclase, 5 wt% olivine, 28 wt% clinopyroxene and 13 wt% orthopyroxene (table 1; figs 4, 5). ilmenite, magnetite and apatite are found in minor amounts totalling about 5 wt% and exhibiting textural relationships similar to those observed for lza. biotite is a rare accessory phase. the gabbronorites are medium grained to coarse grained (3–8 mm) with primocrysts of plagioclase and olivine. the relationships between the two pyroxenes are similar to those observed in lza (fig.  6). the principal difference from lza is that clinopyroxene predominantly appears as granular, inclusion-free grains and is thus defined as a primocryst and the number of oikocrysts is consequently reduced (wager & brown 1967; holness et al. 2015). we also noticed that the modal abundance of clinopyroxene is higher and olivine lower than in lza (table 1; fig. 4). orthopyroxene forms coronas between olivine, plagioclase and feti oxides, most noticeably in the lower 250 m of the subdivision. the top of lzb is easily defined by the transitional appearance of feti oxide-rich layers at the base of lzc. 4.1.1.3 lower (sub)zone c (lzc) the lzc is represented by 160 m of feti oxide gabbronorites (fig. 3) comprising on average 34 wt% plagioclase, 6 wt% olivine, 30 wt% clinopyroxene, 7 wt% orthopyroxene, 17 wt% ilmenite, 7 wt% magnetite and trace amounts of apatite (table 1; figs 4, 5). overall, the rocks are medium grained (1–5 mm), and finer grained than the lzb, with primocrysts of plagioclase, pyroxenes, olivine and feti oxides in a dominating granular texture. the relationships between the two pyroxenes are similar to those observed in lza (fig. 6). reactive coronas of olivine between plagioclase, pyroxene or feti oxides are occasionally present instead of the orthopyroxene coronas of lza to lzb. locally, the feti oxide mode can opx mz lza uzb uza uzc pl cpx + opx ol wt% modal proportions pl cpx lza lzb lzc o px uzc uzb uza lza+b lzc cpx + opx pl m t + il fig. 4 summary of calculated modal mineralogy of the examined gabbros displayed in the triangular (cpx + opx) – pl – ol, opx – pl – cpx (only lz) and (cpx + opx) – (mt + il) – pl diagrams (wt%) with main groupings outlined. cpx: clinopyroxene. pl: plagioclase. ol: olivine. opx: orthopyroxene. mt: magnetite. il: ilmenite. other abbreviations in fig. 2. table 1 average modal makeup of zones and subzones of the ls (wt%) zone pl ol cpx opx il mt ap q kfs uzc 33.3 19.7 33.8 4.6 1.5 5.4 2.6 uzb 41.3 23.5 21.6 6.6 2.7 4.4 uza 43.9 7.3 34.1 3.5 8.3 4.7 0.2 mz 39.7 1.9 33.7 5.9 13.5 5.6 0.1 lzc 34.0 5.5 30.4 7.3 17.1 6.7 0.2 lzb 51.1 4.9 28.0 12.7 2.2 2.3 0.3 lza 56.1 9.4 13.0 18.0 2.2 1.1 1.1 pl: plagioclase. ol: olivine. cpx: high-ca pyroxene. opx: low-ca pyroxene. il: ilmenite. mt: magnetite. ap: apatite; q: quartz. kfs: k-feldspar. complete data set supplied in supplementary file s1. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 13 of 139 geusbulletin.org reach 50 wt% with ilmenite consistently twice as abundant as magnetite (thy et al. 2009a) and occurring as isolated grains or as intergrowths associated with latestage oxidation and exsolution (fig. 6). these lamellar intergrowths occasionally include needles of hercynitic spinel. 4.1.2 middle zone (mz) the mz comprises 260 m of the composite profile. it is an feti oxide gabbronorite (fig. 3) containing on average 40 wt% plagioclase, 2 wt% olivine, 34 wt% clinopyroxene, 6 wt% orthopyroxene, 14 wt% ilmenite and 6 wt% magnetite (table 1; figs 4, 7). again, apatite is a trace mineral (≤0.1 wt%). the mz gabbros are medium grained (2–6 mm) with primocrysts of plagioclase, pyroxenes, olivine and feti oxides and have a granular texture. the textural relationship of clinopyroxene and orthopyroxene is similar to that found in lz (figs 6, 8), while fig. 9 shows a clear antithetical relationship between the abundances of olivine and orthopyroxene. large olivine grains appear abundantly in about 30% of the sampled gabbros throughout the sequence, but particularly in the middle of the mz at 1010–1025 m depth. ilmenite and magnetite intergrowths are similar to those described for the lzc. orthopyroxene coronas are rare and only found at the base of the zone, while olivine coronas are found throughout mz (fig. 8); in both cases, the coronas are texturally similar to those found in the lz. symplectite intergrowths of plagioclase, olivine, and clinopyroxene, as described by holness et al. (2011), are present in the upper part of the mz (fig. 8; 1055–1060 m in the composite profile). platinova resources ltd. obtained several drill cores during 1986–1990, mostly starting in the uz and focusing on the potentials for precious metal deposits as part of the so-called triple group in the upper part of the mz. the drill core through the c. 100 m thick mineralised group in the uppermost part of the mz was not included in this study, although partially covered by surface samples. petrographic and mineralogical details, including the silicates and feti oxides, of this short triple group interval have been documented by andersen (1996, 2006), andersen et al. (1998) and nielsen et al. (2005, 2015, 2019b). 4.1.3 upper zone (uz) the uz varies from gabbronorite in the lower part to gabbro in the upper part, characterised throughout by plagioclase, olivine, orthopyroxene (disappears at 1600 m height), clinopyroxene and feti oxides. the uz represents 987 m of stratigraphy in the composite profile and has been divided into three subzones (uza, uzb and uzc; fig. 3). wager & deer (1939) referred to the uz as “ferro-gabbros” made up of a lower part of “hortonolite” (fa50–70), a middle part of “ferro-hortonolite” (fa70–80) and an upper part of “fayalite” (fa80–100), based on the then common names for the fayalite-content of olivine solid solutions. the uzc was referred to by wager & deer (1939) as the “purple band” because of the purple-brown weathering that coincides with the lzc lza lzb cpx pl ol ox 458205 (592 m) 458211 (7 m) 458287 (875 m) mi mi 2 mm 2 mm 2 mm opx fig. 5 representative petrography of the lz gabbros (samples 458287, 875 m; 458205, 592 m; 458211, 7 m). shown are macrophotographs of typical gabbros for each of the three subzones (lza, b and c). cpx: ca-rich pyroxene (with subordinate intergrown low-ca pyroxene). mi: mica. ol: olivine. opx: low-ca pyroxene (with subordinate intergrown high-ca pyroxene). ox: feti oxides (intergrowth of ilmenite and magnetite). pl: plagioclase. scale bars are 2 mm. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 14 of 139 geusbulletin.org appearance of high iron and manganese pyroxene (ferrobustamite). 4.1.3.1 upper (sub)zone a (uza) the uza makes up 432 m of feti oxide gabbronorite (fig. 3) containing on average 44 wt% plagioclase, 7 wt% olivine, 34 wt% clinopyroxene, 4 wt% orthopyroxene, 8 wt% ilmenite, 5 wt% magnetite and c. 0.2 wt% apatite (table 1; figs 4, 10). the uza gabbronorites are medium-grained (1–5 mm) with primocrysts of plagioclase, pyroxenes, olivine and feti oxides in a mostly granular texture. throughout the subzone, olivine modes are markedly higher than in the mz, while orthopyroxene is relatively rare compared to mz. the textural relationships between the pyroxenes are, however, similar to those observed for mz (figs 8, 11). olivine coronas between plagioclase, pyroxene and feti oxides are present, but rare, while feti oxides are found as isolated grains or intergrown with silicate minerals. 4.1.3.2 upper (sub)zone b (uzb) the uzb consists of 424 m of feti oxide and apatite gabbro (fig. 3) with an average of 40 wt% plagioclase, 24 wt% olivine, 22 wt% clinopyroxene, 7 wt% ilmenite, 3 wt% magnetite, and 4 wt% apatite (table 1; figs 4 and 10). the gabbros are medium grained to coarse grained (2–8 lzb lzb lzc lzalzalza lzalzalza 458208 (703 m) 458278 (710 m) opx cpx il opx cpx pl il 458285 (830 m) pl mt il ol cpx ol 458211 (7 m) ap mi il pl 458216 (125 m) opx ol il pl 458215 (115 m) ap il mi pl il mt pl 458211 (7 m) 458216 (125 m) 458216 (125 m) pl opx cpx ol il mt ol opx opx pl fig. 6 representative pyroxene intergrowths and late crystallising mineralogy and textures in lz gabbros illustrated by backscattered electron (bse) images. samples shown in order of appearance: 458216 (125 m), 458215 (115 m), 458211 (7 m), 458208 (703 m), 458278 (710 m), and 458285 (830 m). ap: apatite. il: ilmenite. mt: magnetite. scale bars are 200 μm or 100 μm, as shown. other abbreviations in fig. 5. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 15 of 139 geusbulletin.org mm) with primocrysts of plagioclase, clinopyroxenes, olivine, feti oxides and apatite in a mostly granular texture. orthopyroxene is occasionally found (e.g. at 1630 m; fig. 11). the most conspicuous features of uzb are the high olivine mode, the near absence of orthopyroxene and the occurrence of abundant apatite, the latter used to define the base of this subzone. up-section, hedenbergite becomes distinctly brown and pleochroic, while olivine is commonly oxidised. coronas of olivine are rare, but otherwise textures are similar to those in uza. 4.1.3.3 upper (sub)zone c (uzc) the uzc comprises 123 m of chiefly ferrograbbros, or more precisely ferrodiorites (fig. 3), with an average of 33 wt% plagioclase, 20 wt% olivine, 34 wt% clinopyroxene, 5 wt% ilmenite, 2 wt% apatite, 5 wt% quartz, and 1 wt% orthoclase (table 1; figs 4, 10). magnetite is scarce or absent. quartz appears as individual grains or in conspicuous granophyric intergrowth with orthoclase, albite and quartz; the latter is thought to have inverted from tridymite (lindsley et al. 1969; larsen & tegner 2006). the base of uzc is defined by a change from granular hedenbergite to a fine-grained interlocking mosaic of green ferrohedenbergite with high manganese and brown ferrohedenbergite with low manganese and high titanium (fig. 12). the green ferrohedenbergite is interpreted as inverted from ferrobustamite, while the brown ferrohedenbergite is thought to have crystallised directly from a late silicate melt (naslund 1984). this mosaic pyroxene texture is shown in fig. 10 (transmitted light) and fig. 11 (backscattered electron (bse) image). the uzc gabbro also contains abundant interstitial granophyre or micropegmatite that increases up-section, reaching up to 14 wt% of the rock composed of a granophyric intergrowth of quartz, albite and orthoclase (figs 13, 14). orthoclase is confined to graphic intergrowths, while quartz and albite also occur as rare granular grains. 4.1.4 liquid immiscibility jakobsen et al. (2005) and jakobsen (2007) presented evidence for liquid immiscibility by examining mineral inclusions in the same set of samples analysed as part of the present study. after homogenising these inclusions, they were able to identify coexisting silica-rich and silica-poor inclusions in apatite and olivine of the uzb and uzc and interpreted these as conjugate immiscible melts. in a more extensive study, jakobsen et al. (2011) showed that melt inclusions trapped in early-crystallised plagioclase could be detected as early as in the upper part of the lz. they suggested that these represented entrapment of variable amounts of two conjugate immiscible melts in an evolving interstitial melt and that the upper part of the ls from about lzc, through mz and uz, crystallised from an emulsion of immiscible melts. 4.2 mineral mode and variation in gabbro density the mineral modes shown in fig. 15 and summarised in table 2 are calculated from bulk-gabbro compositions and the mineral compositions present in the individual gabbro samples. thus, the modes represent solidus or subsolidus conditions and not necessarily liquidus or subliquidus conditions. despite such limitations, several of the features revealed, contribute to our understanding of liquidus conditions. plagioclase is by far the dominant mineral at about 40–60 wt%. the plagioclase content in the lz is constant at about 50 wt% up to lzc, where a fall in plagioclase modes upwards into mz is related to the appearance of feti oxides. plagioclase content is then constant until uzc where it drops followed by a sharp increase towards the top of this zone (fig. 15). the clinopyroxene modes vary between 30 wt% and 50 wt% but are markedly lower in lza (10–20 wt%), where clinopyroxene is interpreted to be interstitial and thus correlated with the amount of trapped melt. furthermore, there is a marked decrease mz 458250 (935 m) pl cpx ox ol 2 mm 2 mm 458251 (953 m) fig. 7 representative petrography of the mz gabbros (samples 458251, 953 m and 458250, 935 m). shown are macrophotographs of typical gabbros. abbreviations in fig. 5. scale bars are 2 mm. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 16 of 139 geusbulletin.org in clinopyroxene modes through uza and into the base of the uzb, replaced by a reversal to increasing modes through the rest of uzb, culminating in a marked increase in uzc followed by a sharp decrease towards the sh. olivine is mostly well below 10 wt% from the lz and until uzb. remarkably, olivine increases up to 25–30 wt% at the base of uzb, followed by a systematic drop further upward into uzb (fig. 15). a minor drop in the upper part of mz is correlated with a corresponding increase in the orthopyroxene content. orthopyroxene shows a marked decline upwards from a maximum of 25 wt% in lza and 5 wt% in the lower part of the mz. in the middle and upper part of the mz, an increase in orthopyroxene is correlated with a drop in olivine. orthopyroxene tapers out in uza and disappears in uzb. both ilmenite and magnetite appear near simultaneously in the base of the lzc and peak in the centre of lzc at about 20 wt% ilmenite and 5 wt% magnetite, despite large modal variations related to layering (thy et al. 2009a; tegner et al. 2009). both oxide modes systematically reduce upwards through the mz and uzb. magnetite is virtually absent in the uzc. apatite shows a strong increase at the base of uzb (holness et al. 2017b) where it first appears as an early mineral with average rocks containing up to c. 8 wt% and subsequently decreases. quartz and orthoclase were only detectable in the uzc, where they sharply increase. there exist few modal data obtained by point counting for skaergaard gabbros (maaløe 1987; conrad & naslund 1989), probably due to the problem of obtaining reliable information from the coarse-grained rocks. most of these data are from maaløe (1974, 1976a, 1978, 1987), pl mz mz 458259 (1010 m) 458263 (1055 m) opx ap pl ilol cpx cpx/ol/pl mz 458258 (1009 m) mt il ol cpx ol pl mz 458262 (1053 m) opx cpx pl mz mz 458249 (930 m) 458254 (988m) pl ol mt il opxopx cpx pl cpx fig. 8 representative pyroxene intergrowths and late crystallising mineralogy and textures in the mz gabbros illustrated by bse images. samples shown in order of appearance: 458249 (930 m), 458254 (988 m), 458262 (1053 m), 458258 (1009 m), 458259 (1010 m) and 458263 (1055 m). abbreviations in figs 5, 6. the pyroxene intergrowth and lamellae are indicated by the grey shading of the images. scale bars are 200 mm or 100 mm, as shown. 0 5 10 15 20 0 2 4 6 8 10 olivine (%) o rt ho py ro xe ne /o liv in e fig. 9 antithetic relationship between the calculated modal contents of olivine and orthopyroxene shown as orthopyroxene/olivine (weight ratio) as a function of olivine (wt%) for observations where the modal content of olivine is above zero. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 17 of 139 geusbulletin.org obtained for the purpose of understanding the origin of layering at selected intervals in the stratigraphy. maaløe (1974, 1987) discussed in detail a short interval (a total of c. 4.7 m) in lza and obtained an average mode of 59 wt% plagioclase, 18 wt% olivine, 14 wt% clinopyroxene, 7 wt% orthopyroxene and 3 wt% feti oxides (on a weight basis). conrad & naslund (1989) examined three 10–15 cm thick, modallyand size-graded, single gabbroic layers in uza that showed a strong anti-correlation between olivine and plagioclase, amounting to a doubling in plagioclase and a complete depletion in olivine modes upward in these layers. taking their data directly from their graphs and recalculating to a weight basis, suggest averages of 41 wt% plagioclase, 32 wt% clinopyroxene, 9 wt% olivine and 18 wt% feti oxides. most of these point-counted modal data for lza and uza compare reasonably well with the calculated modal contents of this study (table 1). knowing the absolute amount of modal orthopyroxene is critical for evaluating potential liquid lines of decent, mineral reactions and the parental-magma composition. it is possible that the apparent high orthopyroxene content in lz may at least in part reflect solidus and subsolidus exsolution from clinopyroxene and may thus not reflect the liquidus content. this may be due to the inability of the electron microprobe analyses to capture the bulk-pyroxene composition despite the broad beam used. holness et al. (2015) optically mapped thin sections, similar to those examined in this study, and obtained average lza compositions of 60 wt% plagioclase, 13 wt% olivine, 18 wt% clinopyroxene, 7 wt% orthopyroxene, and 2 wt% feti oxides. both these optical estimates differ from the estimate using least-squares mixing calculations in this study (table 1), principally by the lower total amount of orthopyroxene as well as a lower proportion of orthopyroxene of the total pyroxenes (0.3 against 0.6). likewise, the cipw normative compositions estimate 58 wt% plagioclase, 15 wt% olivine, 12 wt% clinopyroxene, 10 wt% orthopyroxene and 5 wt% feti oxides for the lza gabbro – closer, but not close enough, to the 56 wt% plagioclase, 9 wt% olivine, 13 wt% clinopyroxene, 18 wt% orthopyroxene and 2 wt% feti oxides calculated by mixing. these differences are no doubt related to the effects of partial un-mixing and exsolution of the pyroxenes and magnetite on the modes calculated based on subsolidus re-equilibrated gabbros. the bulk-gabbro density measured by weighing the dry and clean samples in air and in water reflects the modal variation of the gabbros (fig. 15). the density is constant, around 3.03 ± 0.10 g·cm–3 in the lza and lzb, but markedly increases with the appearance of feti oxides in the lzc to highly variable values of 3.3 to 4.2 g·cm–3 with an average of 3.6 g·cm–3. the uz starts with a slight decrease upward in uza, followed by a reversal to higher values in uzb, reflecting the high iron-rich olivine content in these gabbros, and again followed by a slight decrease in the upper parts of the uzb. 4.3 mineral chemistry 4.3.1 olivine olivine ranges in composition between fo67 and fo01, initially with little systematic upward variation in the lz uzc uzb uza 2 mm 2 mm 2 mm cpx pl ox ol 458653 (2075 m) 90-22-854.5 (1237 m) 90-22-344.1 (1748 m) ap fig. 10 representative petrography of uz gabbros (samples 458653, 2075 m; 90-22-344.1, 1748 m and 90-22-854.5, 1237 m). shown are macrophotographs (transmitted light) of typical gabbros for each of the three subzones (uza, b and c). abbreviations in figs 5, 6. scale bars are 2 mm for all photographs. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 18 of 139 geusbulletin.org (lza fo67.0–fo48.1; lzb fo61.2–fo47.4; lzc fo60.2–fo47.4), followed by a marked and systematic decrease starting in the upper part of the mz (mz fo51.3–fo39.3; uza fo41.7– fo25.1; uzb fo30.2–fo1.8; uzc fo4.5–fo0.8). olivine in the hz reaches fo content of fo66.0 (nwe 1976; holness et al. 2013, 2015) and fo74.3 in the so-called ‘tranquil division’ of lza* of mbs (hoover 1989b). table 3 presents average olivine compositions calculated per formula unit (pfu) of four oxygens. based on 921 analyses acquired in this study, olivine has a near ideal stoichiometry (m2sio4, where m is the total cations mg, fe, mn and ca) with total cations of 3.010 ± 0.014 (1σ) calculated to a formula unit of 4 oxygens (m2.02si0.99o4), si of 0.990 ± 0.014 (1σ), and a strong negative correlation between mg and fe (fe = 2.000 – 0.971 mg; ∑r2 = 0.998) with end points for pure fayalite of 1.999 (fe+mn) and for forsterite of 2.060. manganese systematically increases up-section from 0.011pfu at the base of lza to 0.051 pfu at the base of the uzc and then decreases to 0.010 pfu (fig. 16), because of the appearance of mn-rich ferrobustamite. the average calcium content is 0.10 ± 0.06 (wt% cao) or 0.003 ± 0.002 (ca pfu) with no correlation with mg content. olivine coronas analysed (n = 11) in rocks of lzc to uza (figs 6, 8, 11) are systematically more fo-rich by 1–3 wt% fo compared to the primocryst compositions (fig. 17). the present results are consistent with earlier determinations of olivine compositions made using optical or x-ray methods by analysing mineral separates (deer & wager 1939; wager & deer 1939; yoder & sahama 1957; wager & brown 1967) and later by electron microprobe analyses (e.g. nwe 1976; mcbirney 1989a). 4.3.2 plagioclase plagioclase ranges in composition from an72 to an23 with weak systematic upward decreases within lz (lza an71.6– an51.3; lzb an62.0–an47.9; lzc an57.0–an45.6) and mz (an68.5– an42.1), and continuing with a more marked decrease in uz (uza an47.9–an37.7; uzb an41.1–an30.2; uzc an38.2–an23.4). plagioclase in the lza and lzb may show strong normal zoning, and rarely, reverse zoning (wager & brown 1967; maaløe 1976a; toplis et al. 2008; namur et al. 2014), reflected here in the uncertainties of the drill-core analyses. while plagioclase in the hz reaches similar maximum an contents of an68.9 (maaløe 1976a; holness et al. 2013, 2015), it has been recorded in the mbs to reach an71.4 in the tranquil division of the lza* (hoover 1989b). the hz plagioclase compositions are slightly higher than the maximum values (an69.5) reported from the ubs by salmonsen & tegner (2013). table 4 shows average plagioclase compositions calculated as pfu of uzb 90-22-461.8 (1630 m) cpx opx uzb 90-22-461.8 (1630 m) opx cpx ol pl cpx uzb 90-22-290.8 (1801 m) ap pl ol cpx il uza opx cpx ol mt il pl ol 90-22-809.2 (1282 m) uza 90-22-854.5 (1237 m) opx cpx pl uza 90-22-854.5 (1237 m) opx cpx fig. 11 representative pyroxene intergrowths and late crystallising mineralogy and textures in uza and uzb gabbros illustrated in bse images. samples shown in order of appearance: 90-22-854.5 (1237 m), 90-22-809.2 (1282 m), 90-22-290.8 (1801 m) and 90-22-461.8 (1630 m). abbreviations in figs 5, 6. scale bars are 200 mm or 100 mm, as shown. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 19 of 139 geusbulletin.org eight oxygens. the composition of plagioclase in the granophyres in uzb and uzc is albitic, ranging from an4–7 in interstitial granophyres to an18–27 in pockets of granophyre in pegmatitic gabbro (larsen & brooks 1994; larsen & tegner 2006). it is thus predictable that a whole range of plagioclase compositions from an40 to an0 may occur in the granophyres and associated gabbros. the analyses (n = 959) record near-perfect plagioclase stoichiometry (naalsi3o8–caal2si2o8) with total cations of 5.024 ± 0.017 (1σ) calculated as pfu of 8 oxygens. the major cations all show strong correlation with ca/si = 2.982 – 0.999 ca (∑r2 = 0.983), na+k = 0.972 – 0.894 ca (∑r2 = 0.980), and al = 1.031 + 0.940 ca (∑r2 = 0.975). the minor element ti (0.003) is constant, fe (0.015) shows a systematic increase in uz (tegner 1997), and k (0.025) records an increase starting in mz and continuing into uzb (fig. 18). plagioclase shows dominant oscillatory and normal zoning in the hz and lza (maaløe 1976a). anorthite-rich rims (an55–40) on normally zoned grains have, however, been reported in the lz (toplis et al. 2008; humphreys 2009, 2011; namur et al. 2014). the results of this study can be compared to previous observations using optical methods (wager & deer 1939; carr 1954; gay & muir 1962; wager & brown 1967), analyses of mineral separates (jang & naslund 2001) and electron microprobe analyses (maaløe 1976a; mcbirney 1989a; tegner 1997). 4.3.3 clinopyroxene average analyses of clinopyroxene are shown in table 5. pyroxene varies in compositions from augite to ferroaugite and ferrohedenbergite, without reaching into the salite fields in the pyroxene quadrilateral (poldervaart & hess 1951; fig. 19a). the transmitted light colours of the pyroxenes are transparent to weak brownish, but transition into weak green and further to yellow and violet-brown pleochroic colours in the uzc (wager & deer 1939), where ferrohedenbergite coexists with inverted green ferrobustamite (figs 10, 12). despite extensive exsolution of low-ca pyroxenes in the high-ca pyroxenes, the broad electron-beam technique used in this study has largely succeeded in recording average compositions that only marginally reach into the subcalcic augite fields (fig. 19a). the overall trend seen in the enstatite-wollastonite-ferrosalite (en-wo-fs) triangular diagram is similar to that reported by brown et al. (1957) and brown & vincent (1963) using wet chemical and spectroscopic analyses of mineral separates and by nwe & copley (1975) and nwe (1975, 1976) using electron microprobe analyses. this supports the notion that the average pyroxene compositions largely represent liquidus and subliquidus compositions. most noticeably, the new data show the central keel in the wollastonite variation for medium mg/(mg + fetotal) ratios (fig. 19b), well established for the skaergaard pyroxenes (nwe 1976). pl 458643 (2150 m) ap mn ti bse brown ferrohedenbergite green ferrobustamite green ferrobustamite brown ferrohedenbergite uzc 458643 (2150 m) fig. 12 mineralogy and textures of ferrohedenbergite and inverted ferrobustamite intergrowths in a uzc gabbro (sample 458643, 2150 m) illustrated in a bse density image and elemental-mapped electron kα images (mn and ti). other abbreviations in figs 5, 6. the relationship between the two pyroxenes is seen as green inverted ferrobustamite, characterised by a mosaic of grains, rimmed by a late crystallising brownish ferrohedenbergite. scale bars are 1000 mm. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 20 of 139 geusbulletin.org this peak in the keel coincides with the disappearance of orthopyroxene, except for a couple of occurrences in the uz. the variation in the clinopyroxene compositions parallels that for olivine with initial weak upward decrease in en content (lza en54.0–en38.1; lza en48.3–en36.3; lzc en44.3–en36.9) and a subsequent stronger decrease from the upper mz and into the uz (mz en45.5–en35.1; uza en41.3–en29.4; uzb en32.6–en5.8; uzc en6.6–en0.0). the drop in wollastonite amounts to about 15 wt% from wo39 in lza to wo33 in mz (fig. 19b). augite in the hz reaches a maximum of en41.5 (wo45.0en41.5fs13.5) at about 50 m before the base of the cambridge drill core (nwe 1976). the total amount of cations in clinopyroxenes calculated for 6 oxygens (n = 1060) is 4.023 ± 0.014 (1σ), only slightly over the ideal stoichiometry of 4 cations. the negative correlation between total fe and mg is offset from the ideal 1:1 ratio (fe = 1.123 – 0.829 mg (∑r2 = 0.943)) and does not improve if mn is added to fe (fe + mn = 1.141 – 0.838 mg (∑r2 = 0.942)). this offset is attributed to ferric iron that is calculated as part of the total iron content and may increase with increasing total fe. details of the compositional variation are shown in figs 20 and 21, here normalised to 4 cations and 6 oxygens. aluminium shows a marked decrease as a function of mg from 0.11 pfu throughout the lz and mz and thereafter levels at about 0.05 in uza, forming a sloping plateau until the middle of uzc when it reaches c. 0.04, and followed by a slight decrease to c. 0.03 pfu for uzcsi na k al 458643 (2150 m)458643 (2150 m) 458643 (2150 m) 458643 (2150 m) pl q px ox fig. 13 elemental-mapped electron kα images (si, na, k, al) of interstitial granophyre patches in uzc (sample 458643, 2150 m). intergrown albite (an~0), quartz and orthoclase are seen along margins of trapped plagioclase grain (an~25-30). pl: plagioclase. px: pyroxenes. q: quartz. ox: feti oxides. scale bars are 1000 μm. uzc 458643 (2150 m) ox pl ol q px si fig. 14 textures of melanogranophyre from uzc (458643, 2150 m) illustrated by a bse density image. the sample is composed of granular albitic plagioclase (~an25–30), pyroxenes (see fig. 11), olivine and interstitial granophyre (see fig. 13). pl: plagioclase. px: pyroxenes. ol: olivine. q: quartz. ox: feti oxides (mostly ilmenite). scale bar is 10 mm. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 21 of 139 geusbulletin.org the ferrohedenbergites of the sh (fig. 20). titanium is much lower than al, showing a decrease from c. 0.03 pfu through the lz to 0.02 pfu in mz. in contrast to al, ti forms a slightly increasing plateau from 0.02 to 0.03 pfu, and ends with a terminal decrease from the middle of uzc to c. 0.02 at the sh. aluminium occurs dominantly as tetrahedral aliv with little octahedral alvi and mostly in the uz (not shown). of the analysed clinopyroxenes, only 18% can be calculated to contain octahedral al at an average 0.015 pfu, about a quarter of the average 0.06 pfu content of tetrahedral al. the ti/al ratio increases markedly with decreasing mg (ti/al = 0.929 + 0.761 mg (∑r2 = 0.844)). this translates into a systematic variation as a function of decreasing al from a ti/al of 0.25 in the lza to 0.40 through the lz to mz, followed by a sharp transgression to 1.0 through the uz and capped by a constant ratio in the sh (fig. 21; nwe & copley 1975). manganese shows a systematic increase from 0.01 pfu with a deflection in uza until reaching 0.25 at the base of uzc, followed by a marked decrease to c. 0.01–0.02 pfu in the terminal brown ferrohedenbergites of the sh where it coexists with mn-rich inverted ferrobustamite (fig. 20). sodium is present at about 0.018 ± 0.005 pfu (1σ) through the ls, slightly decreasing past the middle of uzb (not shown). chromium is below the detection limits for the uz but may reach 0.005 pfu in the lz. ferric iron content, estimated by charge-balancing non-quadrilateral elements (alvi+2ti+cr versus na+aliv; papike et al. 1974) is shown in fig. 22. because of the low content of na, cr and alvi in most skaergaard pyroxenes, this is basically a balance of 2ti against aliv. although such estimates do not often correlate well with directly measured values, the results suggest that the fe3+ content of skaergaard pyroxenes is largely constant throughout the lz and mz (0.06–0.04 pfu) and thereafter form a broad up-section increasing plateau towards a small negative content of fe3+ in uzc. the estimated fe3+ shows an initial marked drop in the lz and mz with a subsequently low-sloping trend towards slightly negative values (fig. 22). if fe3+ is estimated by bulk charge balance, the fe3+ content is not markedly different from the non-quadrilateral charge balance, except that the absolute values are somewhat higher (0.05 ± 0.03 pfu, 1σ), with no systematic variation with mg content and no negative values. because of the constancy in fe3+, the fe3+/fe2+ ratio shows a variation with mg content that mimics the fe3+ variation seen in fig. 22 and calculated by the method of papike et al. (1974). if fe3+ cations were expressed as a ratio of the total iron, fe3+/(fe2++fe3+), the results would be relatively constant for the two segments of fig. 22: 0.07 ± 0.04 (1σ) for the uz and 0.28 ± 0.14 for the mz and lz. the latter, however, decreases slightly upward. although caution is warranted using charge-balance calculations, the results suggest that fe3+ (and the fe3+/ fe2+ ratio) decreases through the lz and mz and reaches constant low values, approaching zero fe3+, in the uz. it is, however, possible that the high ti/al ratio correlated with low calculated fe3+ may signal the presence of reduced ti3+ as suggested for some lunar pyroxenes (bence & papike 1972). plagioclase olivine clinopyroxene orthopyroxene ilmenite magnetite apatite quartz orthoclase st ra ti gr ap hi c he ig ht (m ) mode fractions 90 -2 2 lzb lza lzc mz uza uzb uzc 0 0.2 0.4 0.6 0.8 0 400 800 1200 1600 2000 0.1 0.2 0.3 0.4 0.2 0.4 0.6 0.1 0.2 0.3 0.4 0.1 0.2 0.3 0.4 0.1 0.2 0.3 0.4 0.1 0.2 0.1 0.2 0.3 0.4 0.1 0.2 3.0 3.5 density g⋅cm–³ 0 400 800 1200 1600 2000 fig. 15 modal variation of the ls (as weight fractions) as a function of stratigraphic height (m). calculated based on bulk-rock and mineral compositions as described in section 4.2 and table 2. shown are plagioclase, olivine, clinopyroxene, orthopyroxene, ilmenite, magnetite, apatite, quartz and orthoclase. last column is the measured gabbro density (g·cm−3) obtained using archimedes’ principle. vertical red line: location of drill core 90-22. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 22 of 139 geusbulletin.org cation occupancy suggests that the structural tetrahedral site in pyroxene (t-site) is predominantly filled by si and aliv and that the amount of fe3+ occupying octahedral sites (i.e. structural octahedral sites in pyroxene; m-sites) markedly decreases in the lz and mz to low or zero in uz. aluminium occurs predominantly in tetrahedral (aliv) coordination with lesser amounts in octahedral coordination (alvi). the average deficiency in total t-site occupancy (si+alvi) to an ideal value of 2.00 is 0.02 that may be due to fe3+ or other tetrahedral cations that were not analysed. the total of the non-quadrilateral components, calculated according to cawthorn & collerson (1974) and using bulk charge-balance fe3––fe2+ partitioning, amounts to 11–5% that correlates positively with mg. the only non-quadrilateral components are c. 3.5–1% table 2 summary of calculated modes without trapped melt zone stratigraphic height (m) f ol opx cpx pl mt il ap q kfs high low (wt%) (wt%) (wt%) (wt%) (wt%) (wt%) (wt%) (wt%) (wt%) ubs 2169 2165 17.17 22.97 40.80 3.70 1.37 9.10 4.90 uzc 2165 2165 0.000 19.59 27.86 37.79 3.53 0.87 7.06 3.30 uzc 2165 2164 0.000 21.13 32.69 28.92 5.10 1.60 7.10 3.47 uzc 2163 2147 0.001 19.67 32.94 31.55 4.77 1.47 6.84 2.77 uzc 2144 2141 0.003 22.70 36.63 26.27 4.83 1.80 5.50 2.27 uzc 2091 2081 0.010 15.87 42.63 33.40 5.13 1.80 0.83 0.33 uzc 2075 2060 0.012 20.71 40.07 30.44 5.34 1.87 1.00 0.57 uzc 2060 2046 0.014 21.10 34.78 37.68 4.50 1.23 0.44 0.27 uzb 2044 2030 0.016 19.97 24.37 38.97 3.64 1.10 11.35 0.60 uzb 2024 1984 0.019 22.33 30.33 40.07 0.43 4.83 2.00 uzb 1962 1921 0.029 24.02 24.42 42.97 0.17 5.50 2.93 uzb 1902 1862 0.039 22.67 19.17 47.68 0.97 6.00 3.50 uzb 1843 1801 0.050 24.14 19.96 40.86 4.03 6.73 4.28 uzb 1782 1748 0.062 26.20 19.50 41.03 0.67 7.77 4.83 uzb 1728 1691 0.073 27.34 13.36 38.67 3.66 9.37 7.61 uzb 1671 1630 0.085 20.68 21.84 40.55 2.97 7.27 6.70 uzb 1620 1600 0.097 10.86 33.58 40.04 5.54 6.93 3.05 uza 1561 1527 0.112 8.09 4.94 26.94 48.80 3.17 7.93 0.13 uza 1506 1431 0.126 9.11 0.84 32.04 45.95 2.56 9.44 0.07 uza 1403 1343 0.154 6.72 3.72 30.92 46.21 3.80 8.47 0.17 uza 1323 1282 0.177 11.82 32.60 41.79 5.13 8.16 0.50 uza 1268 1237 0.194 7.42 0.07 36.77 41.47 6.85 7.29 0.13 uza 1218 1188 0.210 3.96 1.30 39.66 42.02 3.19 9.66 0.20 uza 1178 1131 0.223 1.57 3.98 37.58 41.22 6.61 8.91 0.13 mz 1098 1055 0.251 0.23 8.31 25.39 46.82 4.49 14.66 0.10 mz 1053 1032 0.267 0.00 8.78 39.12 32.56 0.89 18.65 0.00 mz 1024 1009 0.277 3.47 8.04 35.57 37.25 5.53 10.07 0.07 mz 1003 1002 0.285 0.87 4.95 30.62 42.73 8.12 12.66 0.07 mz 988 961 0.291 4.17 3.93 42.60 27.08 4.01 18.11 0.10 mz 953 930 0.304 2.03 2.89 31.97 37.04 10.79 15.22 0.06 mz 921 900 0.316 1.74 6.13 28.54 49.28 4.48 9.76 0.07 mz 890 875 0.329 3.86 0.03 30.59 42.59 7.08 15.77 0.07 lzc 847 816 0.346 6.15 2.98 27.85 30.17 10.36 21.97 0.53 lzc 808 798 0.362 5.45 2.74 23.74 31.74 6.88 29.23 0.22 lzc 784 742 0.372 2.08 7.44 43.12 34.19 4.49 8.59 0.10 lzc 723 703 0.398 6.61 8.73 34.26 46.49 1.78 2.00 0.13 lzb 703 681 0.406 4.86 11.94 27.45 50.94 1.51 2.60 0.69 lzb 634 580 0.437 3.49 8.82 33.36 51.78 0.40 1.76 0.40 lzb 558 488 0.472 9.12 11.72 27.68 49.11 0.26 1.95 0.16 lzb 447 367 0.526 6.54 15.07 23.37 52.53 0.45 1.90 0.13 lzb 346 221 0.576 2.76 14.72 25.17 50.39 4.55 2.31 0.11 lzb 177 173 0.667 3.60 16.72 16.80 60.00 0.38 2.30 0.19 lza 161 137 0.676 7.95 16.59 14.57 57.37 1.20 2.03 0.30 lza 125 107 0.696 10.84 22.63 10.84 51.39 1.40 2.57 0.33 lza 96 27 0.712 13.54 16.99 13.66 52.70 0.66 2.13 0.31 lza 7 0.764 7.11 17.22 18.32 53.65 1.30 2.10 0.30 calculated modes normalised to a total gabbro mode of 100%. each determination is based on three samples, except at 7 m where only one sample was used. compare to table 24 for the same modes given with preferred trapped-melt content. the high quartz content at the top of uzb is caused by a single sample characterised as being a melanogranophyre. f: liquid remaining after tegner et al. (2009), (eq.) 1. complete data set in supplementary file s1. ol: olivine. opx: orthopyroxene. cpx: clinopyroxene. pl: plagioclase. mt: magnitite. ap: apetite. q: quartz. kfs: orthoclase. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 23 of 139 geusbulletin.org ca-ti-tschermakite (catial2o6), c. 4–0% ca-fe3+-tschermakite (cafe3+alsio6), 2–0% ca-tschermakite (caal2sio6) and 3–1.5% jadeite (naalsi2o6). the calculated molecules indicate a final propagated error of silica excess amounting to 0.03 ± 0.03 pfu (1σ). most calculated and total non-quadrilateral molecules show a systematic decrease throughout the lz to uza and are replaced by a constant or slight increase throughout uzb to uzc – all as a positive function of decreasing mg (fig. 23). jadeite decreases through the lz to uza and is constant at about 1.5% throughout uzb to uzc. the ca-ti-tschermakite varies in the lz to uza, and mirrors jadeite, but shows a systematic increase in uzb, followed by a drop in uzc. the calculated ca-ferri-tschermakite molecule is dependent on the fe3+–fe2+ partitioning model chosen. the model used here results in a systematic decrease table 3 average olivine compositions zone stratigraphic height (m) n sio2 (wt%) tio2 (wt%) feo (wt%) mno (wt%) mgo (wt%) cao (wt%) nio (wt%) total (wt%) fo high low ave. sd ave. sd ave. sd ave. sd ave. sd ave. sd ave. sd ave. sd ubs 2169 2165 28 28.91 0.248 0.06 0.015 69.50 0.583 0.65 0.091 0.32 0.075 0.36 0.429 0.03 0.021 99.83 0.01 0.002 uzc 2165 2165 13 29.35 0.188 0.09 0.014 70.14 0.737 0.48 0.051 0.04 0.019 0.16 0.175 0.03 0.013 100.28 0.00 0.015 uzc 2165 2164 14 29.87 0.215 70.59 0.421 0.50 0.031 0.09 0.103 0.02 0.019 101.07 0.00 0.003 uzc 2163 2147 23 29.30 0.193 70.12 0.615 0.79 0.080 0.06 0.038 0.02 0.013 100.29 0.00 0.001 uzc 2144 2141 14 28.99 0.439 0.10 0.015 69.77 0.823 1.07 0.095 0.08 0.018 0.14 0.091 0.04 0.019 100.18 0.00 0.000 uzc 2091 2081 20 29.14 0.215 68.42 0.503 1.44 0.071 0.44 0.083 0.03 0.013 99.47 0.01 0.002 uzc 2075 2061 13 29.11 0.268 66.93 0.486 1.64 0.091 0.56 0.063 0.02 0.012 98.27 0.01 0.002 uzc 2060 2046 28 29.37 0.258 67.13 0.617 1.74 0.152 1.03 0.101 0.02 0.017 99.29 0.03 0.003 uzb 2044 2030 15 29.61 0.896 0.09 0.025 68.23 1.957 1.13 0.064 0.19 0.049 0.16 0.058 0.03 0.024 99.44 0.01 0.002 uzb 2024 1984 26 29.89 0.409 0.09 0.010 66.32 0.831 1.37 0.081 2.26 0.111 0.22 0.284 0.02 0.012 100.16 0.06 0.003 uzb 1962 1921 36 30.10 0.353 0.08 0.010 64.64 1.044 1.22 0.068 3.74 0.260 0.09 0.050 0.02 0.015 99.89 0.09 0.006 uzb 1902 1862 24 30.56 0.273 62.12 0.404 1.11 0.041 5.39 0.230 0.02 0.013 99.20 0.13 0.005 uzb 1843 1801 22 30.93 0.269 0.08 0.018 60.96 1.005 0.94 0.063 7.58 0.318 0.08 0.086 100.56 0.18 0.008 uzb 1782 1748 18 31.25 0.312 0.06 58.57 0.590 0.90 0.034 9.47 0.342 0.10 100.34 0.22 0.007 uzb 1728 1691 17 32.07 0.195 55.25 0.562 0.85 0.089 11.57 0.176 0.03 0.017 99.77 0.27 0.003 uzb 1671 1630 60 32.17 0.261 54.22 0.312 0.81 0.079 12.65 0.255 0.04 99.89 0.29 0.005 uzb 1620 1600 18 31.95 0.288 55.74 0.404 0.84 0.059 11.10 0.285 0.01 99.65 0.26 0.005 uza 1561 1527 27 32.16 0.326 54.58 0.482 0.83 0.057 12.31 0.200 99.89 0.29 0.004 uza 1506 1431 21 32.36 0.234 0.03 0.021 54.01 0.685 0.79 0.029 13.50 0.260 0.07 0.055 0.05 100.81 0.31 0.006 uza 1403 1343 23 32.68 0.319 0.04 0.012 51.74 0.594 0.72 0.041 15.41 0.213 0.10 0.036 0.03 100.72 0.35 0.005 uza 1323 1282 25 33.20 0.204 48.74 0.453 0.73 0.064 17.17 0.289 0.02 0.009 99.86 0.39 0.005 uza 1268 1237 25 33.18 0.340 0.09 49.53 0.602 0.71 0.069 16.86 0.290 0.08 0.018 100.45 0.38 0.005 uza 1218 1188 8 32.85 0.227 49.51 0.411 0.69 0.018 16.65 0.388 0.09 0.019 99.80 0.37 0.006 uza 1178 1131 15 33.04 0.280 0.06 48.32 0.747 0.67 0.036 17.67 0.258 0.06 0.09 0.029 99.91 0.39 0.006 mz 1098 1055 3 33.70 0.211 0.06 0.038 47.12 0.653 0.52 0.059 20.21 0.327 0.13 0.052 0.04 0.056 101.78 0.43 0.002 mz 1053 1032 mz 1024 1009 16 33.42 0.335 0.03 0.022 46.90 0.741 0.59 0.058 20.82 0.249 0.09 0.026 0.06 0.057 101.91 0.44 0.005 mz 1003 1002 mz 988 961 21 34.28 0.318 0.05 0.023 43.90 0.705 0.56 0.068 23.28 0.568 0.09 0.027 0.10 0.07 102.27 0.49 0.010 mz 953 930 mz 921 900 8 33.51 0.782 0.16 0.021 46.03 1.057 0.55 0.071 21.90 0.071 0.06 0.020 0.05 0.08 102.25 0.46 0.006 mz 890 875 5 34.34 0.258 0.03 0.025 40.61 1.158 0.52 0.072 23.21 0.916 0.09 0.047 0.10 0.08 98.89 0.50 0.007 lzc 847 816 15 34.92 0.246 0.05 0.026 39.39 1.053 0.50 0.096 25.89 0.618 0.09 0.041 0.03 0.06 100.88 0.54 0.011 lzc 808 798 12 35.27 0.195 0.04 0.034 38.88 0.680 0.44 0.080 26.84 0.466 0.10 0.028 0.10 0.08 101.68 0.55 0.007 lzc 784 742 16 34.77 0.359 0.03 0.020 40.89 0.845 0.47 0.069 25.23 0.315 0.08 0.029 0.08 0.07 101.54 0.52 0.008 lzc 723 703 13 34.64 0.233 0.04 0.029 41.22 0.554 0.48 0.103 25.27 0.309 0.12 0.017 0.09 0.09 101.86 0.52 0.004 lzb 703 681 17 34.40 0.219 0.05 0.049 42.23 0.446 0.48 0.105 25.20 0.266 0.11 0.032 0.07 0.07 102.52 0.52 0.005 lzb 634 580 11 34.35 0.289 0.06 0.025 41.73 0.567 0.52 0.079 23.24 0.168 0.13 0.036 0.08 0.07 100.11 0.50 0.004 lzb 558 488 31 34.82 0.306 0.04 0.025 41.11 0.438 0.50 0.075 25.37 0.379 0.11 0.033 0.08 0.05 102.02 0.52 0.005 lzb 447 367 13 34.70 0.316 0.03 0.023 40.96 0.558 0.43 0.065 26.28 0.338 0.12 0.030 0.10 0.08 102.61 0.53 0.008 lzb 346 221 20 35.31 0.279 0.04 0.027 37.27 0.691 0.43 0.042 27.69 0.444 0.09 0.021 0.09 0.07 100.93 0.57 0.004 lzb 177 173 16 34.63 0.278 0.04 0.030 40.15 0.863 0.51 0.106 25.41 0.269 0.10 0.042 0.08 0.07 100.93 0.53 0.006 lza 161 137 18 35.28 0.378 0.04 0.015 37.14 0.796 0.46 0.080 27.16 0.280 0.09 0.029 0.08 0.07 100.23 0.57 0.008 lza 125 96 19 35.32 0.261 0.03 0.017 36.29 0.645 0.47 0.039 28.15 0.339 0.07 0.026 0.14 0.06 100.48 0.58 0.005 lza 96 27 21 35.37 0.428 0.03 0.025 36.67 1.637 0.48 0.078 27.96 1.131 0.07 0.021 0.10 0.06 100.69 0.57 0.025 lza 7 6 35.78 0.502 0.04 0.021 34.05 2.735 0.44 0.071 29.99 2.086 0.07 0.023 0.09 0.09 100.45 0.61 0.025 n: number of individual mineral analyses used for each calculated average (ave.) for three consecutive samples with 1σ sd where n > 2. stratigraphic height is given as high and low referring to the interval used for averaging. complete data set in supplementary file s1. fo as mole fractions. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 24 of 139 geusbulletin.org throughout the lz to uza and extending into the lower part of the uzb, after which fe3+ can no longer be calculated. finally, ca-tschermakite likewise decreases throughout the lz to uza after which octahedral al can no longer be assigned. these variations mimic the overall elemental variations in pyroxenes by showing a noticeable break in compositional variation between the lz to uza and uzb to uzc. most noticeable is the suggestion that fe3+ occurs in tetrahedral coordination and that octahedral al is not suggested to be present in uzb to uzc. general cation substitutions for terrestrial augites are dominated by single exchange components (fe2+,mn)(mg)–1, ca(mg)–1, fe3+(al)–1, and the coupled exchange components tial2(mgsi2)–1, fe3+al(mgsi)–1 and naal(2mg)–1 (cawthorn & collerson 1974; robinson 1980; papike 1980; basaltic volcanism study project 1981; sack & ghiorso 1994). in some respects, these substitutions differ from the observed uppermost uz pyroxenes. the differences stem from the generally low si content of hedenbergite, requiring most al to be allocated to the tetrahedral site and largely prohibiting octahedral al as well as octahedral fe3+. the pearson matrices of table 6 show that the dominating non-quadrilateral correlations in the overall intrusion are aliv–fe3+ and ti–aliv. the main difference is between the two stratigraphic parts, such that in the upper part (uzb to uzc) the quadrilateral correlation (mg,fe2+)–ca becomes important. in summary, the dominating pattern of cation variation is a negative correlation between fe2+ and mg with a positive correlation between fe2+ and ca, becoming noticeable in the uz. the tetrahedral site is occupied by si and aliv, while fe3+ occurs in octahedral sites together with a small amount of alvi. the fe3+ calculated by charge balancing decreases in the lz and mz and reaches low concentrations in uz. although the absolute variation of ti decreases upward in the stratigraphy, the ti/al ratios show an overall increase upward in the stratigraphy, controlled by corresponding decreasing al. the observation that al largely occupies tetrahedral sites in the skaergaard clinopyroxenes concurs with previous work (muir 1951; brown et al. 1957; brown & vincent 1963; nwe 1975, 1976; nwe & copley 1975). direct measurements of the ferric iron content of the skaergaard augite to hedenbergite series by brown et al. (1957) and brown & vincent (1963) indicate a marked decrease in the average fe3+/fe2+ ratio from the lz and 0.00 0.01 0.02 0.03 0.04 0.05 0.06 0.07 0.00 0.20 0.40 0.60 0.80 1.00 1.20 1.40 m n pf u (4 o xy ge n) mg pfu (4 oxygen) olivine lzmzuzauzb uzc corona ave. & sd all fig. 16 olivine mineral chemistry showing mn as a function of mg as cations per formula unit (pfu) normalised to 4 oxygens (table 3). the approximate locations of the zone boundaries of the ls are indicated along the top of the graph. small blue dots: analysed point compositions. black dots: average compositions for individual thin sections with 1σ standard deviation (sd). open circles: corona compositions shown in fig. 17. ave.: average. 30 35 40 45 50 55 60 65 30 35 40 45 50 55 60 65 fo (mol%) primocryst fo (m ol % ) c or on a 1:1 fig. 17 composition of olivine coronas (fo mol%, see fig. 6) as a function of average primocryst composition (fo mol%, shown with the 1:1 relationship (dashed line). the linear equation (solid line) is foc = 1.160 fop – 5.589 (∑r2 = 0.990, n = 11), where c is corona and p is primocryst. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 25 of 139 geusbulletin.org mz (0.120) into the uz (0.049), thus supporting our observation of fe3+ content in the uz pyroxenes. 4.3.4 ferrobustamite under the general heading of ‘ferrobustamite’ is described the fine mosaic of green hedenbergite that characterises the pyroxenes of the uzc and is believed to have inverted from a triclinic ferrobustamite (naslund 1984). bustamite is a triclinic pyroxenoid, (mn,ca,fe)2si2o6, that is similar to wollastonite with fe and mn substituting for ca (1.5% mn, 55.3% fe2+ and 43.2% ca). thus, the skaergaard ferrobustamite was traditionally referred to as ferrowollastonite with fe substituting for ca to approximately equal amounts (wager & deer 1939; brown & vincent 1963; lindsley et al. 1969; nwe & copley 1975). however, in many table 4 average plagioclase compositions zone stratigraphic height (m) n sio2 (wt%) tio2 (wt%) al2o3 (wt%) feo (wt%) mgo (wt%) cao (wt%) na2o (wt%) k2o (wt%) total (wt%) an high low ave. sd ave. sd ave. sd ave. sd ave. sd ave. sd ave. sd ave. sd ave. sd ubs 2169 2165 15 61.43 0.496 0.03 0.020 23.96 0.392 0.36 0.081 0.01 0.004 5.22 0.314 8.54 0.217 0.44 0.093 99.99 0.247 0.016 uzc 2165 2165 15 61.28 0.327 0.04 0.016 24.08 0.235 0.44 0.031 0.01 0.004 5.70 0.227 8.08 0.206 0.56 0.051 100.19 0.272 0.012 uzc 2164 2164 23 60.96 0.479 0.03 0.007 24.01 0.211 0.45 0.161 0.02 0.004 5.93 0.162 7.76 0.202 0.46 0.049 99.62 0.288 0.009 uzc 2163 2147 26 60.92 0.537 0.04 0.014 24.02 0.339 0.43 0.054 5.73 0.277 7.96 0.174 0.54 0.071 99.63 0.274 0.014 uzc 2144 2141 18 60.86 0.474 0.05 0.018 24.32 0.210 0.45 0.050 0.02 0.003 6.11 0.251 7.79 0.143 0.55 0.051 100.15 0.292 0.012 uzc 2091 2081 20 59.75 0.474 24.20 0.358 0.45 0.049 6.64 0.364 7.24 0.159 0.47 0.103 98.75 0.325 0.016 uzc 2075 2061 21 59.13 0.433 24.21 0.194 0.43 0.030 6.63 0.225 7.35 0.170 0.41 0.059 98.17 0.323 0.012 uzc 2060 2046 9 58.90 0.541 24.45 0.278 0.44 0.032 0.01 0.006 6.96 0.253 7.07 0.231 0.41 0.032 98.24 0.342 0.012 uzb 2044 2030 16 59.21 0.442 0.04 0.020 24.82 0.190 0.45 0.036 0.02 0.006 6.47 0.221 7.75 0.171 0.48 0.036 99.25 0.306 0.011 uzb 2024 1984 27 59.30 0.476 0.05 24.79 0.318 0.44 0.027 0.01 0.004 6.91 0.273 6.95 0.158 0.39 0.029 98.85 0.344 0.012 uzb 1962 1921 29 58.58 0.568 0.04 24.87 0.344 0.44 0.025 0.04 0.015 7.08 0.331 6.89 0.271 0.46 0.036 98.40 0.349 0.015 uzb 1902 1862 27 57.82 0.758 25.07 0.316 0.39 0.045 7.22 0.220 6.79 0.143 0.50 0.053 97.78 0.357 0.010 uzb 1843 1801 18 58.36 0.381 25.36 0.152 0.40 0.060 7.50 0.159 6.60 0.181 0.48 0.031 98.71 0.372 0.010 uzb 1782 1748 18 57.66 0.476 25.41 0.183 0.36 0.070 7.49 0.154 6.73 0.159 0.56 0.056 98.22 0.365 0.009 uzb 1728 1691 18 58.15 0.388 25.46 0.176 0.41 0.044 7.70 0.202 6.53 0.144 0.46 0.026 98.71 0.381 0.011 uzb 1671 1630 43 57.54 0.500 25.63 0.215 0.35 0.048 7.74 0.165 6.53 0.131 0.47 0.034 98.26 0.382 0.009 uzb 1620 1600 17 57.73 0.392 25.90 0.264 0.33 0.078 7.92 0.192 6.33 0.194 0.42 0.027 98.63 0.396 0.010 uza 1561 1527 18 56.51 0.259 26.13 0.196 0.34 0.041 8.34 0.150 6.34 0.184 0.44 0.029 98.10 0.407 0.010 uza 1506 1431 17 56.30 0.437 26.40 0.234 0.31 0.048 8.63 0.256 6.27 0.176 0.43 0.034 98.33 0.419 0.013 uza 1403 1343 19 55.97 0.619 0.05 26.69 0.247 0.28 0.039 8.91 0.301 6.10 0.223 0.40 0.052 98.39 0.434 0.014 uza 1323 1282 19 55.63 0.520 26.76 0.341 0.28 0.054 8.90 0.324 6.07 0.156 0.39 0.034 98.03 0.435 0.015 uza 1268 1237 9 55.34 0.892 26.82 0.346 0.27 0.026 9.02 0.382 6.02 0.288 0.44 0.049 97.90 0.439 0.021 uza 1218 1188 26 55.33 0.814 26.70 0.387 0.26 0.027 8.89 0.413 5.89 0.187 0.44 0.039 97.50 0.440 0.017 uza 1178 1131 24 55.54 0.675 0.09 0.027 26.96 0.336 0.30 0.057 0.03 0.013 9.15 0.265 6.05 0.197 0.40 0.023 98.52 0.443 0.014 mz 1098 1055 16 55.62 0.642 0.09 0.020 27.73 0.314 0.38 0.142 0.03 0.017 9.63 0.310 6.25 0.197 0.35 0.034 100.09 0.451 0.012 mz 1053 1032 12 54.33 0.638 0.11 0.025 27.57 0.381 0.41 0.097 0.03 0.022 9.92 0.398 6.16 0.163 0.28 0.061 98.81 0.464 0.017 mz 1024 1009 23 54.78 0.532 0.10 0.027 27.72 0.367 0.37 0.139 0.04 0.028 9.98 0.218 6.00 0.152 0.34 0.032 99.34 0.470 0.011 mz 1003 1002 23 54.60 0.926 0.09 0.031 28.07 0.364 0.35 0.073 0.04 0.024 10.18 0.351 5.98 0.234 0.32 0.031 99.62 0.476 0.016 mz 988 961 24 55.15 0.506 0.11 0.018 27.88 0.526 0.38 0.062 0.04 0.017 10.28 0.323 5.85 0.214 0.35 0.034 100.03 0.483 0.016 mz 953 930 25 55.14 0.597 0.11 0.029 28.08 0.389 0.33 0.068 0.04 0.018 10.26 0.406 5.89 0.286 0.32 0.037 100.18 0.482 0.020 mz 921 900 24 53.62 0.872 0.10 0.032 28.44 0.657 0.39 0.135 0.05 0.050 10.94 0.719 5.44 0.372 0.24 0.048 99.22 0.519 0.036 mz 890 875 19 54.28 0.843 0.12 0.035 27.70 0.594 0.45 0.104 0.07 0.050 10.52 0.411 5.66 0.206 0.26 0.028 99.05 0.500 0.016 lzc 847 816 20 54.44 0.373 0.11 0.029 27.65 0.409 0.42 0.102 0.09 0.095 10.54 0.345 5.51 0.184 0.29 0.088 99.06 0.505 0.017 lzc 808 798 19 54.58 0.598 0.11 0.027 28.47 0.492 0.57 0.429 0.06 0.059 10.57 0.402 5.69 0.201 0.26 0.052 100.30 0.499 0.015 lzc 784 742 20 53.95 0.542 0.10 0.030 28.55 0.459 0.46 0.254 0.06 0.061 10.85 0.395 5.60 0.212 0.30 0.097 99.89 0.508 0.017 lzc 723 703 18 53.85 0.478 0.12 0.045 28.75 0.336 0.40 0.102 0.04 0.017 11.15 0.301 5.21 0.211 0.32 0.026 99.85 0.532 0.014 lzb 703 681 20 53.21 0.827 0.10 0.021 28.72 0.310 0.44 0.099 0.04 0.014 11.01 0.337 5.45 0.182 0.28 0.034 99.24 0.519 0.015 lzb 634 580 18 53.30 0.469 0.10 0.023 28.96 0.374 0.36 0.069 0.05 0.019 11.54 0.346 5.05 0.200 0.31 0.041 99.68 0.548 0.017 lzb 558 488 29 53.01 0.837 0.11 0.019 29.03 0.414 0.44 0.116 0.06 0.057 11.56 0.550 4.96 0.336 0.31 0.054 99.45 0.553 0.028 lzb 447 367 20 52.48 0.849 0.10 0.027 29.61 0.555 0.46 0.131 0.06 0.057 11.99 0.640 4.93 0.267 0.30 0.059 99.93 0.564 0.024 lzb 346 221 18 51.88 0.526 0.10 0.034 30.03 0.375 0.41 0.096 0.05 0.025 12.52 0.332 4.66 0.219 0.26 0.038 99.91 0.589 0.017 lzb 177 173 19 51.68 0.503 0.09 0.032 30.36 0.396 0.40 0.097 0.04 0.036 12.79 0.336 4.51 0.192 0.25 0.082 100.12 0.602 0.015 lza 161 137 17 52.41 1.098 0.10 0.024 30.13 0.731 0.37 0.050 0.04 0.027 12.40 0.835 4.77 0.435 0.27 0.083 100.49 0.581 0.040 lza 125 96 18 51.73 0.854 0.10 0.022 30.54 0.581 0.40 0.122 0.05 0.062 12.82 0.656 4.50 0.398 0.25 0.073 100.40 0.603 0.033 lza 96 27 18 51.81 1.564 0.08 0.029 30.01 0.968 0.31 0.070 0.03 0.024 12.66 1.061 4.50 0.638 0.26 0.052 99.66 0.599 0.054 lza 7 6 51.42 1.162 0.10 0.013 30.58 0.716 0.33 0.026 0.06 0.017 12.72 0.916 4.48 0.534 0.21 0.101 99.90 0.604 0.047 n: number of individual mineral analyses used for each calculated average (ave.) for three consecutive samples with 1σ sd where n > 2. stratigraphic height is given as high and low referring to the interval used for averaging. complete data set in supplementary file s1. an as mole fractions. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 26 of 139 geusbulletin.org respects, the skaergaard inverted bustamite differs from classic occurrences of bustamite in metasomatised skarn deposits by their higher fe2+ and lower mn (mason 1975; deer et al. 1978). bustamite inverts to a fine-grained mixture of wollastonite and ferrohedenbergite or to pure ferrohedenbergite (brown & guy 1960) during cooling around 950°c for compositions and conditions relevant to the uzc of the skaergaard intrusion (lindsley et al. 1969; rutstein 1971; rutstein & white 1971; deer et al. 1978). the green ferrohedenbergite is easily identified not only petrographically (fig. 10), but also compositionally, by low ti and high mn compared to the coexisting brown ferrohedenbergite (fig. 12; table 7). the mosaic grains of the green ferrohedenbergite are first detected intermittently just below the uzc at 2051 m in the stratigraphy, become abundant in uzc and persist into the sh. the green ferrohedenbergite shows little or no compositional variation as a function of mg content or stratigraphic position. the fe2+ content is constant at 1.084 ± 0.048 and fe3+ at 0.014 ± 0.011 pfu (n = 289), suggesting a small, but generally persistent, ferric iron content in contrast to the coexisting brown ferrohedenbergite (fig. 22), where the presence of fe3+ is not detected by charge-balance calculations. the mg content of the green ferrohedenbergite is 0.021 ± 0.017, mn 0.030 ± 0.011, al 0.009 ± 0.002, ti 0.003 ± 0.002 and na 0.012 ± 0.007 pfu identifying hedenbergite inverted from bustamite by low ti (and al and ti/ al) and relatively high mn compared to the coexisting brown ferrohedenbergite (figs 20, 21). the corresponding concentrations for the coexisting brown ferrohedenbergite (n = 123) are mn 0.014 ± 0.006, al 0.030 ± 0.007, ti 0.027 ± 0.007 and na 0.016 ± 0.006 pfu. the mn/fe2+ ratio for the green ferrohedenbergite (0.028 ± 0.011) is double that for the coexisting brown ferrohedenbergite (0.013 ± 0.006 pfu), further supporting the contention that the former is a polymorph of mn-enriched ferrobustamite (mason 1975). another significant feature of the green ferrohedenbergite is that the tetrahedral sites (si+al+fe3+) are filled (t-sites = 2.000 ± 0.014 pfu), leaving some al and fe3+ in octahedral site positions. the result is that they show low non-quadrilateral components with low ca-ti-tschermakite and high ca pfu (8 oxygen) fe p fu (8 o xy ge n) 0 0.01 0.02 0.03 0.04 0.2 0.3 0.4 0.5 0.6 0.7 k pf u (8 o xy ge n) ave. & sd plagioclase 0 0.01 0.02 0.03 0.04 0.05 lzmzuzauzbuzc fig. 18 plagioclase mineral chemistry showing k and fe as a function of ca as cations per formula unit (pfu) normalised to 8 oxygens (table 4). approximate locations of the ls zone boundaries are indicated along the top of the graph. small blue dots: analysed point compositions. black dots: average compositions for individual thin sections with 1σ sd. ave.: average. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 27 of 139 geusbulletin.org ca-ferri-tschermakite when compared to the coexisting brown ferrohedenbergites. 4.3.5 orthopyroxene orthopyroxene appears throughout the lz, mz and into uza in variable amounts and compositions within the intermediate (inverted) pigeonite and enstatite-ferrosilite fields of poldervaart & hess (1951; fig. 19a). the analyses (n = 127) suggest subsolidus equilibration from clinopyroxene (including pigeonite) for a subset of the analyses by exsolution and grain-coarsening resulting in a decrease in the wollastonite content to a low of 2% into the enstatite-ferrosilite field. on the other hand, a subset of analyses is consistent with the overall variation (table 8) expected for inverted pigeonite with a wollastonite content of c. 10% and similar to values reported by brown et al. (1957) using wet chemical analyses of mineral separations and nwe (1976) using electron microprobe analyses. table 5 average clinopyroxene compositions zone   stratigraphic height (m) n sio2 (wt%) tio2 (wt%) al2o3 (wt%) feo (wt%) mno (wt%) mgo (wt%) cao (wt%) na2o (wt%) cr2o3 (wt%) total (wt%)   mg# high low ave. sd ave. sd ave. sd ave. sd ave. sd ave. sd ave. sd ave. sd ave. sd ave. sd ubs 2169 2165 22 46.36 0.420 0.96 0.108 0.74 0.129 31.33 0.550 0.30 0.050 0.85 0.300 18.65 0.420 0.19 0.041 0.02 0.006 99.40 0.05 0.015 uzc 2165 2165 15 46.84 0.310 0.84 0.244 0.59 0.113 32.33 0.370 0.24 0.060 0.03 0.020 19.02 0.380 0.22 0.094 0.02 0.002 100.12 0.00 0.001 uzc 2165 2164 18 46.57 0.430 0.87 0.201 0.59 0.117 32.33 0.570 0.23 0.040 0.02 0.010 19.16 0.310 0.23 0.109 0.02 0.007 100.02 0.00 0.000 uzc 2163 2147 21 46.73 0.310 0.75 0.217 0.51 0.119 32.12 0.440 0.31 0.110 0.05 0.020 18.97 0.320 0.21 0.011 0.02 0.006 99.69 0.00 0.002 uzc 2144 2141 15 46.62 0.340 0.75 0.087 0.55 0.041 32.27 0.330 0.48 0.040 0.11 0.040 18.80 0.140 0.22 0.031 0.01 0.004 99.81 0.01 0.002 uzc 2091 2081 14 45.61 0.120 0.86 0.109 0.62 0.051 31.39 0.340 0.60 0.090 0.39 0.070 18.70 0.220 0.18 0.013 0.01 98.35 0.02 0.004 uzc 2075 2061 114 46.74 0.250 0.57 0.073 0.47 0.063 30.10 0.650 0.78 0.060 0.65 0.080 18.62 0.290 0.18 0.018 0.01 0.001 98.10 0.04 0.004 uzc 2060 2046 35 47.02 0.220 0.90 0.129 0.66 0.071 29.37 0.210 0.73 0.070 1.39 0.090 18.68 0.150 0.18 0.014 0.01 98.93 0.08 0.005 uzb 2044 2030 17 46.69 0.290 0.88 0.058 0.68 0.062 31.12 0.130 0.63 0.030 1.05 0.200 18.99 0.230 0.29 0.076 0.01 0.002 100.34 0.06 0.010 uzb 2024 1984 28 47.31 0.430 1.03 0.026 0.80 0.121 28.13 0.510 0.61 0.060 2.74 0.210 18.21 0.310 0.21 0.058 99.04 0.15 0.012 uzb 1962 1921 31 47.20 0.610 0.98 0.030 0.82 0.093 26.13 0.540 0.54 0.040 4.35 0.290 18.23 0.380 0.20 0.024 0.01 98.46 0.23 0.014 uzb 1902 1862 25 47.54 0.460 0.95 0.031 0.90 0.083 23.24 0.400 0.49 0.050 6.01 0.190 17.86 0.260 0.22 0.057 97.21 0.32 0.007 uzb 1843 1801 20 48.29 0.320 0.90 0.023 0.92 0.070 22.49 0.370 0.47 0.040 7.34 0.220 18.05 0.510 0.23 0.028 0.01 0.002 98.71 0.37 0.009 uzb 1782 1748 17 48.98 0.230 0.89 0.033 1.01 0.083 21.02 0.230 0.46 0.030 9.10 0.220 17.65 0.380 0.23 0.040 99.34 0.44 0.007 uzb 1728 1691 17 48.38 0.740 0.84 0.023 0.98 0.087 20.11 0.500 0.48 0.060 9.49 0.190 17.18 0.270 0.23 0.010 97.68 0.46 0.007 uzb 1671 1630 10 49.28 0.290 0.81 0.059 1.04 0.076 19.21 0.340 0.43 0.050 10.45 0.160 16.80 0.240 0.24 0.055 98.27 0.49 0.006 uzb 1620 1600 20 49.54 0.410 0.84 0.038 1.09 0.109 19.38 0.550 0.46 0.050 10.66 0.160 16.89 0.840 0.23 0.017 0.01 99.09 0.50 0.009 uza 1561 1527 17 49.55 0.250 0.87 0.040 1.08 0.088 19.48 0.370 0.45 0.030 10.78 0.220 16.83 0.170 0.23 0.031 99.26 0.50 0.008 uza 1506 1431 28 49.61 0.410 0.79 0.132 1.16 0.099 19.16 1.260 0.43 0.070 10.99 0.240 16.69 1.170 0.22 0.043 0.02 0.016 99.08 0.51 0.016 uza 1403 1343 25 49.77 0.420 0.74 0.129 1.24 0.117 18.01 0.730 0.44 0.060 11.80 0.220 16.79 0.660 0.20 0.045 0.02 0.014 99.01 0.54 0.009 uza 1323 1282 21 49.74 0.290 0.68 0.051 1.21 0.122 17.10 0.650 0.42 0.070 12.23 0.180 16.50 0.720 0.21 0.031 0.01 0.002 98.08 0.56 0.009 uza 1268 1237 11 49.62 0.550 0.63 0.085 1.33 0.174 17.20 0.870 0.42 0.100 12.24 0.380 16.36 1.000 0.22 0.045 98.01 0.56 0.011 uza 1218 1188 11 50.00 0.300 0.61 0.062 1.24 0.096 18.31 0.910 0.42 0.100 12.48 0.250 15.73 1.010 0.20 0.010 0.01 99.00 0.55 0.010 uza 1178 1131 26 50.00 0.390 0.71 0.079 1.33 0.167 17.88 1.570 0.40 0.070 13.02 0.420 15.31 1.620 0.21 0.051 0.04 0.021 98.89 0.57 0.016 mz 1098 1055 23 49.83 0.380 0.91 0.265 1.54 0.122 17.06 0.690 0.32 0.050 12.97 0.210 17.06 0.500 0.26 0.041 0.03 0.008 99.98 0.58 0.011 mz 1053 1032 24 49.75 0.360 0.82 0.156 1.52 0.098 15.98 0.900 0.39 0.060 13.56 0.300 17.18 0.790 0.24 0.049 0.04 0.019 99.47 0.60 0.012 mz 1024 1009 24 49.82 0.490 0.77 0.076 1.58 0.198 16.01 1.700 0.35 0.060 13.31 0.570 17.30 1.930 0.27 0.113 0.03 0.017 99.45 0.60 0.018 mz 1003 1002 28 49.56 0.610 0.93 0.186 1.57 0.108 16.75 0.960 0.38 0.070 13.35 0.320 17.13 1.000 0.23 0.045 0.03 0.012 99.92 0.59 0.011 mz 988 961 26 50.15 0.510 0.88 0.112 1.62 0.134 16.83 1.240 0.36 0.060 13.59 0.260 16.71 1.390 0.25 0.039 0.04 0.026 100.44 0.59 0.016 mz 953 930 21 50.22 0.320 0.88 0.097 1.70 0.140 15.01 0.840 0.37 0.080 13.64 0.290 17.86 0.930 0.23 0.033 0.04 0.026 99.96 0.62 0.010 mz 921 900 29 49.71 0.510 0.86 0.197 1.70 0.121 15.85 0.920 0.34 0.090 13.73 0.210 17.18 0.690 0.20 0.046 0.03 0.023 99.61 0.61 0.008 mz 890 875 21 50.38 0.430 0.88 0.096 1.79 0.123 14.76 0.690 0.33 0.050 14.00 0.340 17.45 0.720 0.24 0.064 0.03 0.031 99.85 0.63 0.008 lzc 847 816 25 50.19 0.320 0.95 0.128 1.81 0.152 14.12 1.190 0.28 0.090 13.87 0.470 18.10 1.610 0.25 0.041 0.04 0.014 99.61 0.64 0.012 lzc 808 798 22 50.39 0.430 1.01 0.129 1.90 0.138 13.89 0.700 0.30 0.050 14.09 0.380 18.00 0.900 0.29 0.052 0.03 0.022 99.91 0.64 0.009 lzc 784 742 22 49.93 0.760 0.91 0.232 1.82 0.167 13.30 1.070 0.25 0.060 14.01 0.540 18.78 1.260 0.27 0.058 0.07 0.029 99.34 0.65 0.015 lzc 723 703 21 50.34 0.300 0.94 0.176 1.93 0.325 13.41 1.080 0.26 0.060 14.03 0.430 18.54 1.210 0.28 0.074 0.05 0.032 99.77 0.65 0.013 lzb 703 681 24 49.65 0.480 0.91 0.111 1.95 0.169 13.65 0.710 0.29 0.070 13.97 0.340 18.36 0.760 0.27 0.053 0.03 0.016 99.06 0.646 0.011 lzb 634 580 23 49.84 0.640 1.00 0.111 1.97 0.249 13.19 1.360 0.29 0.070 13.99 0.560 17.95 1.480 0.27 0.053 0.08 0.043 98.59 0.654 0.019 lzb 558 488 30 50.33 0.450 0.89 0.223 1.77 0.316 13.33 0.950 0.30 0.080 14.03 0.380 18.32 1.160 0.28 0.055 0.06 0.028 99.31 0.652 0.014 lzb 447 367 19 49.91 0.450 0.96 0.122 2.00 0.198 13.40 1.110 0.23 0.090 14.24 0.500 18.25 1.090 0.28 0.051 0.09 0.024 99.36 0.655 0.016 lzb 346 221 17 50.17 0.520 0.92 0.167 2.09 0.182 12.09 1.210 0.23 0.070 14.68 0.640 18.69 1.690 0.29 0.060 0.13 0.052 99.30 0.684 0.017 lzb 177 173 16 50.22 0.670 0.88 0.087 2.15 0.262 12.03 1.860 0.30 0.070 14.32 0.900 19.25 2.110 0.28 0.052 0.13 0.070 99.57 0.680 0.026 lza 161 137 16 50.70 0.270 0.81 0.152 2.19 0.235 11.04 1.150 0.27 0.070 14.75 0.710 19.20 1.490 0.31 0.057 0.09 0.081 99.36 0.704 0.018 lza 125 96 17 50.35 0.520 0.83 0.158 2.13 0.189 10.75 0.630 0.24 0.070 14.66 0.650 19.76 1.330 0.35 0.069 0.06 0.057 99.15 0.709 0.007 lza 96 27 18 50.21 0.870 0.88 0.270 2.08 0.282 10.95 1.830 0.27 0.060 15.09 0.970 19.19 2.060 0.30 0.070 0.07 0.041 99.02 0.710 0.026 lza 7 4 50.14 0.230 0.87 0.083 2.09 0.201 11.08 2.130 0.19 0.080 15.67 0.380 18.17 1.580 0.26 0.056 0.16 0.126 98.63 0.716 0.042 n: number of individual mineral analyses used for each calculated average (ave.) for three consecutive samples with 1σ sd where n > 2. stratigraphic height is given as high and low referring to the interval used for averaging. complete data set in supplementary file s1. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 28 of 139 geusbulletin.org like the augite-ferrohedenbergites, the low-ca pyroxenes have tetrahedral sites that are mostly filled by al and fe3+ (t-sites = 1.989 ± 0.024 pfu) with total fe3+ estimated by charge balance as 0.075 ± 0.004 pfu. the minor elements ti, al and na are all below c. 0.05 pfu giving a total non-quadrilateral component of 2.7% ± 1.0% made-up of jadeite (0.36% ± 0.42%), ca ferri-tschermakite (0.87% ± 0.42%), ca-ti -tschermakite (0.90% ± 0.47%) and ca-tschermakite (0.60% ± 0.48%; fig. 23). the propagated error of silica excess is 0.041 ± 0.039 pfu (1σ). 4.3.6 feti oxides feti oxides appear throughout the ls, although first in lzc in appreciable amounts as ilmenite and magnetite primocrysts. the highest amounts of both phases are found in the centre of lzc (fig. 15). ilmenite is often homogeneous without exsolution or oxidation lamella of hematite. magnetite on the other hand is composed mostly of an intergrowth of magnetite and ilmenite, resulting from exsolution of ulvöspinel lamella and subsequent oxidation to secondary ilmenite (vincent 1960). ilmenite compositions are calculated based on a formula camg cafe en fs mz uza uzb uzc lz ferrohedenbergite magnesium pigeonite intermediate pigeonite ferriferous pigeonite en di op sid e ferrosalite diopside subcalcic ferroaugite ferroaugite hedenbergite clinoenstatite clinoferrosilite wollastonite subcalcic augite salite augite camg cafe en fs a b fig. 19 the pyroxene quadrilateral diagram for the skaergaard pyroxenes on a molecular basis. a: nomenclature of poldervaart & hess (1951) with the cloud of analyses from this study. b: details of analysed pyroxenes (table 5). small solid grey dots: individual analyses. small open green circles: inverted ferrobustamite of the uzc. large open circles: average analyses for individual sections. thin grey lines: tie-lines between coexisting average high-ca and low-ca pyroxenes. iron was calculated as total fe. black dashed lines: solidus trends for the skaergaard intrusion established by nwe (1976). en: enstatite. fs: ferrosilite. camg: diopside. cafe: hedenbergite. approximate locations of the skaergaard zone boundaries are indicated along the high-ca pyroxene solidus trend. abbreviations for ls divisions in fig. 2. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 29 of 139 geusbulletin.org 0 0.02 0.04 0.06 0.08 0.10 0.12 0 0.01 0.02 0.03 0.04 0.05 0.06 0 0.01 0.02 0.03 0.04 0.05 0.06 0.0 0.2 0.4 0.6 0.8 1.0 0 0.2 0.4 0.6 0.8 1.0 1.2 1.4 a l p fu (4 c at io ns , 6 o xy ge ns ) ferrobustamite sh augite-ferrohedenbergite ferrobustamite sh ti p fu (4 c at io ns , 6 o xy ge ns ) sh ferrobustamite m n pf u (4 c at io ns , 6 o xy ge ns ) ferrobustamite sh lz mz uza uzb uzc mg pfu (4 cations, 6 oxygens) to ta l f e pf u (4 c at io ns , 6 o xy ge ns ) fig. 20 variation diagrams as a function of mg content for selected cations of augite-ferrohedenbergite pyroxenes calculated per formula unit (pfu) of 4 cations and 6 oxygens. iron has been partitioned between fe2+ and fe3+ using charge balance (papike et al. 1974). blue dots: total database of high-ca pyroxenes. small green dots: inverted ferrobustamite (or green ferrohedenbergite) of uzc. black dots: average composition for individual thin sections with 1σ sd. approximate locations of the ls zone boundaries are shown in the total fe (bottom) panel. sh (sandwich horizon) refers to the brown ferrohedenbergite of the uzc. abbreviations for ls divisions in fig. 2. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 30 of 139 geusbulletin.org 0.0 0.2 0.4 0.6 0.8 1.0 1.2 1.4 0.0 0.2 0.4 0.6 0.8 1.0 0.00 0.01 0.02 0.03 0.04 0.05 0.06 0.00 0.02 0.04 0.06 0.08 0.10 0.12 0.0 0.2 0.4 0.6 0.8 1.0 0 0.1 0.2 0.0 0.2 0.4 0.6 0.8 1.0 0.0 0.2 0.4 0.6 0.8 1.0 1.2 1.4 1/2 1/4 1/1 1/3 mg pfu (4 cations, 6 oxygens) ti p fu (4 c at io ns , 6 o xy ge ns ) al pfu (4 cations, 6 oxygens) sh ferrobustamite ca pfu (4 cations, 6 oxygens) m g pf u (4 c at io ns , 6 o xy ge ns ) ferrobustamite sh augite-ferrohedenbergite low-ca pyroxene ti /a l ( ca ti on s) ferrobustamite mg pfu (4 cations, 6 oxygens) iv a l p fu (4 c at io ns , 6 o xy ge ns ) ferrobustamite augite-ferrohedenbergite augite-ferrohedenbergite augite-ferrohedenbergite mz uz lz mz uza uzb uzc fig. 21 variation diagrams as a function of mg, ca, and al content for elected cations and ratios of augite-ferrohedenbergite pyroxene calculated per formula unit (pfu) of 4 cations and 6 oxygens. iron has been partitioned between fe2+ and fe3+ using charge balance (papike et al. 1974). symbols and notations in fig. 20. mg vs. ca (top panel) shows both clinoand orthopyroxenes (cf. fig. 19). https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 31 of 139 geusbulletin.org unit of 2 cations and 3 oxygens, while magnetite is calculated based on 3 cations and 4 oxygens. iron is partitioned between fe2+ and fe3+ based on charge balance. the solid-solution components are calculated as hematite and ilmenite for ilmenite and ulvöspinel and magnetite for magnetite. representative coexisting magnetite and ilmenite are shown in table 9. ilmenite (n = 405) contains an average of 8% hematite (fig. 24a) and appreciable amounts of mn (0.012 ± 0.006 pfu) and mg (0.054 ± 0.039 pfu). magnetite (n = 248) generally appears in small amounts throughout the ls together with ilmenite. the magnetite content in the solid solution varies considerably reaching much higher values (fig. 24a) than those found experimentally on skaergaard-like melts (thy et al. 2008); presumably reflecting exsolution and re-equilibration in the natural material (vincent & phillips 1954; vincent 1960; buddington & lindsley 1964). of the minor elements, only al (0.118 ± 0.057 pfu) and mg (0.029 ± 0.023 pfu) occur in appreciable amounts in the magnetite solid solution (table 9). vanadium occurs in both ilmenite (v2o3 < 0.3%) and magnetite (v2o3 < 1.7%; vincent & phillips 1954; jang & naslund 2003; jang et al. 2001) but was not analysed during this study. the t (°c) and oxygen fugacity normalised to the nickel-nickel oxide (nno) oxygen buffer and expressed as ∆logfo2 (nno) for the coexisting magnetite and ilmenite pairs were estimated using the geothermometry-oxybarometry technique of ghiorso & evans (2008) and shown in table 9. 4.3.7 apatite apatite first occurs in uzb (fig. 11) as an early crystallising primocryst. throughout the ls, apatite appears as a later interstitial phase (nash 1976; holness et al. 2011; pedersen et al. 2021) and in late-forming pegmatitic veins (nash 1976; brown & peckett 1977; larsen 1992; sonnenthal 1992; larsen & brooks 1994). the analysed apatite (n = 164) is fundamentally composed of ca (9.693 pfu) and p (5.767 pfu) with small amounts of si (0.044 pfu), fe (0.057 pfu) and mn (0.006 pfu) calculated to 26 anions (o,f,cl,oh; table 10; fig. 24b). thus, the calculated average formulae ca9.67p5.81o24(f,cl,oh)2 is relatively close to ideal apatite of ca10p6o24(f,cl,oh)2 and not dissimilar to the data of nash 0.00 0.02 0.04 0.06 0.08 0.10 0.12 0.14 0.16 na + aliv a lvi + 2 ti + c r mz–lzuz fe3+ = 0.00 bust. 0.00 0.05 0.10 0.15 fe3+ = 0.04–0.06 fe3+ = 0.10–0.20 augite-ferrohedenbergite –0.10 –0.05 0.00 0.05 0.10 0.15 fe 3+ p fu (4 c at io ns , 6 o xy ge ns ) ferrobustamite mg pfu (4 cations, 6 oxygens) 0.0 0.2 0.4 0.6 0.8 1.0 mz uz fig. 22 ferric iron estimate for the high-ca pyroxenes using the charge balance equation fe3+ = aliv + na – alvi – cr – 2ti of papike et al. (1974). the fe3+ content decreases systematically from 0.10–0.20 pfu in the uzc green inverted ferrobustamite (bust.), to 0.04–0.06 in the lz and mz augite, and decreases further in uz augite-ferrohedenbergite to reach negative charge balance values. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 32 of 139 geusbulletin.org (1976) and brown & peckett (1977). the measured fand cl-monovalent anions and the balance-calculated oh in the apatite structure vary as a function of stratigraphic position. a mixture of (f+cl+oh) occurs in the lz, mz and uza. in uzb, (f+oh) occurs where apatite becomes an early-forming mineral phase, and in uzc it becomes nearly dry (only f remaining; fig. 24b). this variation in the anion content indicates the exsolution of a cl-rich vapour phase as suggested by sonnenthal (1992) and mcbirney (1995) at the base of the uz and a second oh-vapour boiling in uzc as evident at lower levels and also in studies of pegmatitic veins (larsen & brooks 1994). further detailed studies are, however, needed to unravel these implications beyond the present study of the near liquidus relations (cf. pedersen et al. 2021). 4.3.8 biotite mica appears as a late crystallising phase in the lza and lzb in interstitial patches often coexisting with feti oxides and apatite. this mica is a biotite, intermediate between phlogopite and annite with mg/(mg+fe2+) ratios between 0.72 and 0.52 (table 11; fig. 24c) and with low contents of alvi and fe3+, both assumed to fill the t-sites (calculated to 22 oxygens). although the contents of f and cl were not analysed, nash (1976) determined the monovalent anions in the skaergaard biotite to be dominantly hydroxyl with f/(f+oh) ratios of c. 0.1 and little or no cl. 4.4 cryptic mineral variations the silicate and feti oxide primocrysts show systematic compositional variation as a function of stratigraphic position through the ls (table 12), most noticeably in the mg/(mg+fe) ratios of the mafic minerals and the ca/ (ca+na) ratio of plagioclase. this ‘hidden’ compositional variation has been referred to as cryptic variation or layering (wager & deer 1939; wager & brown 1967; irvine 1982; mcbirney 1989a). cryptic variation has widely mg pfu (4 cations, 6 oxygens)mg pfu (4 cations, 6 oxygens) ca2si2o6 (wollastonite) non-quadrilateral caalalsio6 (ca-tschermak) catialalo6 (ca-ti-tschermak) cafe3+fe3+sio6 (ca-ferri-tschermak) ferrobustamite orthopyroxene augite-ferrohedengite naalsi2o6 (jadeite) % m ol ec ul ar c om po ne nt 0 0.2 0.4 0.6 0.8 1.0 1.2 0 10 20 30 40 50 0 0.2 0.4 0.6 0.8 1.0 1.2 0 3 6 9 12 15 0 1 2 3 4 5 0 1 2 3 4 5 0 1 2 3 4 5 0 1 2 3 4 5 fig. 23 percentage of the main modelled non-quadrilateral molecular components of the pyroxene solid solution as a function of mg pfu. shown are both the high-ca (solid grey circles) and the low-ca pyroxenes (open circles). see section 4.3.3 for a discussion of the choice of molecular components. the non-quadrilateral component is the total of all components, excluding wollastonite, enstatite and ferrosilite. ferrobustamite is the green ferrohedenbergite believed to have inverted from ferrobustamite (small green circles). the only quadrilateral component shown is wollastonite. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 33 of 139 geusbulletin.org been seen as related to variations in liquidus temperature and parental-magma compositions resulting from fractional crystallisation or other types of compositional changes in a closed or partially open magma chamber. corresponding cryptic mineral trends are observed in the mbs and ubs (e.g. salmonsen & tegner 2013). 4.4.1 olivine the olivine composition shows a systematic upward increase in iron from fo61.1±3.6 at the base of lza to fo0.8±0.2 in uzc and the sh (table 12; fig. 25). the upward change in fo content varies systematically from about 2.0% per 100 m of stratigraphy in the lz to uza, to 7.3% at the table 6 pearson product moment correlation coefficient (r) matrices for non–quadrilateral and quadrilateral cations in clinopyrozene subzones uzc–uzb ti alvi aliv fe3+ fe2+ mn mg ca na ti 1.00 alvi –0.17 1.00 aliv 0.39 –0.46 1.00 fe3+ –0.23 –0.51 0.91 1.00 fe2+ 0.10 –0.39 –0.45 –0.23 1.00 mn –0.01 0.06 –0.05 –0.07 –0.03 1.00 mg –0.11 0.37 0.47 0.25 –1 –0.01 1.00 ca 0.15 –0.48 –0.34 –0.16 0.92 –0.02 –0.94 1.00 na –0.11 –0.02 0.08 0.46 –0.14 –0.06 0.13 –0.16 1.00 zones lz–uza ti alvi aliv fe3+ fe2+ mn mg ca na ti 1.00 alvi –0.23 1.00 aliv 0.40 –0.6 1.00 fe3+ –0.04 –0.7 0.87 1.00 fe2+ –0.13 0.29 –0.83 –0.76 1.00 mn –0.19 0.26 –0.64 –0.54 0.72 1.00 mg 0.15 –0.32 0.69 0.60 –0.7 –0.55 1.00 ca 0.09 –0.2 0.51 0.49 –0.72 –0.52 0.05 1.00 na 0.23 –0.09 0.44 0.45 –0.49 –0.31 0.22 0.38 1.00 table 7 average ferrobustamite compositions per sample zone sample id stratigraphic height n sio2 (wt%) tio2 (wt%) al2o3 (wt%) feo (wt%) mno (wt%) mgo (wt%) cao (wt%) na2o (wt%) cr2o3 (wt%) total (wt%) mg# (m) ave. sd ave. sd ave. sd ave. sd ave. sd ave. sd ave. sd ave. sd ave. sd ave. sd ubs 458634 2168 13 47.39 0.127 0.12 0.034 0.26 0.121 31.24 1.097 0.37 0.120 0.09 0.037 19.99 1.075 0.21 0.072 0.02 0.008 99.63 0.000 0.003 uzc 458636 2165 9 47.37 0.123 0.12 0.045 0.33 0.076 31.67 1.141 0.48 0.093 0.05 0.033 19.67 0.986 0.27 0.055 0.01 0.006 99.96 0.003 0.002 uzc 458637 2165 4 46.61 0.462 0.10 0.040 0.23 0.105 32.30 1.904 0.39 0.076 0.01 0.003 19.80 1.321 0.35 0.134 0.02 0.005 99.79 0.000 0.001 uzc 458638 2165 9 47.53 0.153 0.13 0.018 0.28 0.110 31.83 1.245 0.34 0.030 0.04 0.012 19.60 1.217 0.18 0.043 0.02 0.000 99.95 0.002 0.001 uzc 458639 2165 8 47.59 0.623 0.14 0.066 0.26 0.107 31.52 1.244 0.37 0.054 0.05 0.023 19.70 1.173 0.14 0.036 0.02 0.002 99.81 0.003 0.002 uzc 458640 2164 4 46.92 0.373 0.18 0.047 0.28 0.066 32.83 0.986 0.40 0.021 0.03 0.011 18.78 0.387 0.42 0.399 0.01 0.000 99.85 0.002 0.001 uzc 458641 2164 5 47.70 0.135 0.11 0.055 0.22 0.087 31.10 0.666 0.44 0.124 0.08 0.037 20.10 0.692 0.15 0.028 0.02 0.008 99.92 0.005 0.003 uzc 458642 2162 1 46.86 0.09 0.38 34.05 0.46 0.01 17.36 0.27 0.02 99.51 0.001 uzc 458643 2150 9 47.70 0.366 0.17 0.093 0.21 0.109 31.41 0.865 0.51 0.061 0.06 0.035 19.93 0.805 0.23 0.040 0.01 0.005 100.22 0.003 0.002 uzc 458644 2147 7 47.50 0.235 0.19 0.090 0.26 0.104 31.79 1.538 0.64 0.171 0.08 0.062 19.15 1.302 0.14 0.032 0.01 0.003 99.77 0.005 0.004 uzc 458645 2144 9 46.96 0.199 0.14 0.051 0.35 0.180 32.43 1.348 0.84 0.088 0.09 0.069 18.25 0.942 0.24 0.084 0.01 0.001 99.31 0.005 0.004 uzc 458646 2144 9 47.90 0.263 0.08 0.081 0.16 0.102 30.05 0.704 0.73 0.084 0.11 0.072 20.40 0.800 0.24 0.054 0.01 0.000 99.69 0.006 0.004 uzc 458647 2141 1 46.39 0.263 0.14 0.000 0.85 0.000 32.60 0.000 0.57 0.000 0.16 0.000 18.73 0.000 0.38 0.000 99.81 0.009 uzc 0.7 2091 116 47.14 0.613 0.07 0.035 0.21 0.126 32.63 0.899 1.10 0.117 0.44 0.080 18.21 0.568 0.10 0.039 0.02 0.010 99.92 0.023 0.004 uzc 458652 2089 ? 46.69 0.369 0.06 0.072 0.12 0.096 29.59 1.026 1.06 0.132 0.36 0.064 20.53 0.666 0.08 0.039 0.02 0.005 98.34 0.021 0.003 uzc 458653 2075 9 46.86 0.162 0.04 0.016 0.15 0.041 28.96 1.015 1.15 0.123 0.45 0.112 20.28 0.574 0.08 0.013 0.01 0.006 97.99 0.027 0.007 uzc 20.8 2071 102 46.73 0.533 0.05 0.035 0.08 0.086 30.39 0.616 1.04 0.101 0.61 0.092 18.15 0.763 0.11 0.047 0.01 0.000 97.18 0.035 0.005 uzc 31.1 2061 9 46.29 0.561 0.09 0.045 0.16 0.223 30.04 0.976 1.10 0.116 0.61 0.045 17.88 0.975 0.13 0.073 96.30 0.035 0.002 uzc 458658 2060 1 46.57 0.561 0.11 0.000 0.19 0.000 34.65 0.000 0.74 0.000 0.23 0.025 17.05 0.000 0.32 0.000 99.85 0.012 uzc 40.9 2051 12 47.63 0.507 0.09 0.075 0.12 0.149 28.70 1.046 1.25 0.150 0.80 0.087 19.73 0.791 0.08 0.040 98.39 0.047 0.005 n: number of individual mineral analyses used for each calculated average (ave.) for three consecutive samples with 1σ sd where n > 2. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 34 of 139 geusbulletin.org top of uzb. the variation compares to predicted ideal fractional crystallisation trends in a closed basaltic system (maaløe 1976b; morse 1996). the noticeably high variability in the lza and lzb decreases in the mz and particularly in the uz is attributable to elevated residual melt content in the lower part of the intrusion (tegner et al. 2009). a couple of irregularities may be seen first within the lzc as a regression towards higher forsterite when feti oxides become dominant, and then again at the uza to uzb boundary where modal olivine markedly increases. nwe (1976) and holness et al. (2015) reported olivine compositions in the hz up to fo65, similar to the maximum value noted by mcbirney (1989a) in his eastern section. the compositional variation of ubs olivine reverses the ls variation from fo0 at the sh to fo57 at the top of lz’ (naslund 1984). the only minor element in olivine that shows a systematic variation with stratigraphic height is manganese, which modestly increases from the base of lza until the lower part of uzb. it then markedly increases to the base of uzc where it reaches 0.25 pfu, and then decreases to 0.01 pfu at the top of uzc to the same level as the base of lza (fig. 25). the average calcium content is 0.10 ± 0.06 (wt% cao) or 0.003 ± 0.002 (ca pfu) and shows no detectable variation with stratigraphic height, as also observed by nwe (1976). the average calcium content of the skaergaard olivine equals the maximum limit of cao for plutonic olivine reported by simkin & smith (1970). 4.4.2 plagioclase the cryptic variation is most noticeable in plagioclase primocrysts that vary from calcic at the base of lza (an63.2±5.7) to sodic at the sh of uzc (an24.7±1.6; table 12) defining a near linear variation as a function of stratigraphic height (fig. 26) with a few deviations at the base of the mz. an unusually high average primocryst an content of 61 was found in a single sample from the lower part of mz (900 m) that may reflect erroneous sampling of a block from the basaltic roof of the intrusion (mcbirney 1989a; irvine table 8 average orthopyroxene compositions zone stratigraphic height (m) n sio2 (wt%) tio2 (wt%) al2o3 (wt%) feo (wt%) mno (wt%) mgo (wt%) cao (wt%) na2o (wt%) cr2o3 (wt%) total (wt%) mg# high low ave. sd ave. sd ave. sd ave. sd ave. sd ave. sd ave. sd ave. sd ave. sd ave. sd uzb 1671 1630 4 49.37 0.082 0.21 0.103 0.30 0.157 33.09 2.189 0.67 0.050 12.92 0.064 2.77 1.974 0.04 0.019 0.00 0.000 99.37 0.410 0.016 uzb 1620 1600 1 49.86 0.22 0.39 37.37 0.76 10.64 2.67 0.06 0.00 101.96 0.337 uza 1561 1527 uza 1506 1431 uza 1403 1343 1 49.79 0.15 0.41 32.60 0.67 14.25 2.31 0.02 0.00 100.19 0.438 uza 1323 1282 4 51.23 0.105 0.20 0.051 0.48 0.090 31.68 0.725 0.65 0.058 15.48 0.385 1.84 0.660 0.08 0.107 0.01 0.001 101.64 0.465 0.009 uza 1268 1237 2 50.91 0.34 0.58 28.63 0.64 16.24 2.74 0.03 0.01 100.13 0.503 uza 1218 1188 3 50.64 0.295 0.13 0.024 0.62 0.035 30.69 0.808 0.67 0.059 15.81 0.683 1.93 0.792 0.02 0.007 0.00 0.000 100.52 0.479 0.014 uza 1178 1131 5 50.41 0.235 0.25 0.022 0.51 0.076 30.81 1.243 0.63 0.082 15.78 0.454 2.00 0.467 0.02 0.016 0.00 0.000 100.40 0.477 0.016 mz 1098 1055 6 50.82 0.249 0.29 0.079 0.65 0.073 27.26 1.395 0.50 0.120 16.42 0.919 4.59 1.592 0.04 0.037 0.03 0.000 100.58 0.518 0.015 mz 1053 1032 6 50.29 0.695 0.28 0.116 0.73 0.167 26.19 2.113 0.57 0.106 17.88 1.068 3.50 1.890 0.10 0.119 0.01 0.014 99.55 0.549 0.022 mz 1024 1009 8 50.11 1.148 0.44 0.669 0.78 0.088 26.14 0.880 0.53 0.128 17.19 0.912 4.59 1.582 0.09 0.123 0.01 0.000 99.87 0.541 0.011 mz 1003 1002 4 49.75 0.879 0.43 0.285 0.79 0.142 26.84 2.654 0.49 0.103 16.84 0.504 4.96 2.437 0.06 0.043 0.01 0.000 100.18 0.528 0.019 mz 988 961 6 50.97 0.192 0.29 0.038 0.73 0.074 26.39 1.250 0.50 0.068 18.02 0.278 3.82 1.036 0.03 0.038 0.02 0.011 100.77 0.549 0.012 mz 953 930 6 51.10 0.545 0.33 0.041 0.67 0.098 26.71 0.393 0.52 0.106 18.32 0.896 2.85 1.510 0.03 0.020 0.05 0.018 100.59 0.550 0.010 mz 921 900 1 51.20 0.36 0.55 26.22 0.50 18.20 3.75 0.08 0.02 100.88 0.553 mz 890 875 lzc 847 816 1 51.23 0.35 2.04 22.82 0.40 19.05 4.08 0.00 0.03 99.99 0.598 lzc 808 798 1 50.96 0.38 0.72 25.37 0.48 20.00 2.08 0.00 0.00 100.00 0.584 lzc 784 742 3 51.21 0.574 0.33 0.029 0.83 0.127 23.38 0.721 0.40 0.110 20.30 0.606 3.27 1.597 0.08 0.042 0.07 0.016 99.87 0.607 0.006 lzc 723 703 2 52.16 0.35 0.64 25.56 0.56 19.79 2.78 0.06 0.00 101.88 0.580 lzb 703 681 1 50.66 0.34 0.78 25.12 0.48 18.99 3.93 0.03 0.00 100.33 0.574 lzb 634 580 lzb 558 488 2 51.63 0.39 0.85 22.80 0.48 20.65 3.08 0.06 0.02 99.96 0.617 lzb 447 367 2 51.16 0.35 1.00 25.36 0.39 19.71 2.46 0.05 0.05 100.53 0.580 lzb 346 221 1 51.25 0.36 0.87 21.78 0.42 20.81 3.31 0.05 0.02 98.86 0.629 lzb 177 173 5 51.41 0.284 0.29 0.030 0.95 0.188 24.01 1.610 0.48 0.105 20.91 1.122 1.94 1.823 0.02 0.015 0.02 0.017 100.03 0.608 0.023 lza 161 137 4 52.47 0.691 0.31 0.029 0.84 0.055 21.57 0.345 0.41 0.095 22.35 0.675 1.85 0.663 0.01 0.034 0.03 0.044 99.83 0.649 0.009 lza 125 96 5 52.22 0.385 0.33 0.061 0.90 0.089 20.13 0.672 0.35 0.042 23.51 0.693 1.36 0.172 0.03 0.024 0.04 0.019 98.87 0.587 0.013 lza 96 27 2 52.08 0.31 0.97 19.56 0.37 23.82 1.31 0.03 0.06 98.51 0.68 lza 7 2 51.82 0.31 0.92 22.30 0.40 22.66 1.12 0.01 0.01 99.54 0.644 n: number of individual mineral analyses used for each calculated average (ave.) for three consecutive samples with 1σ sd where n > 2. stratigraphic height is given as high and low referring to the interval used for averaging. complete data set in supplementary file s1. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 35 of 139 geusbulletin.org ta bl e 9 re pr es en ta tiv e co ex is tin g m ag ne tit e an d ilm en ite c om po si tio ns p er s am pl e m ag ne tit e zo ne   sa m pl e id st ra ti g ra ph ic he ig ht n si o 2 (w t% ) ti o 2 (w t% ) al 2o 3 (w t% ) fe o * (w t% ) m no (w t% ) m go (w t% ) ca o (w t% ) cr 2o 3 (w t% ) n io (w t% ) fe o (w t% )   fe 2o 3 (w t% )   to ta l (w t% )   x u lv   x m t   t (° c)   ∆l og fo 2 (n n o ) (m ) av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd u zc 45 86 36 21 65 3 0. 23 0. 12 3 15 .6 8 2. 68 2 0. 56 0. 10 2 79 .9 6 1. 92 9 0. 34 0. 09 7 0. 01 0. 00 0 0. 02 0. 02 0. 03 0. 00 6 0. 05 0. 01 9 45 .7 9 37 .9 7 10 0. 66 0. 46 1 0. 53 9 68 2 –1 .5 8 u zc 45 86 37 21 64 1 1. 72 4. 36 0. 49 88 .6 4 0. 12 0. 01 0. 01 0. 01 0. 03 37 .6 2 56 .6 9 10 1. 05 0. 13 6 0. 86 4 50 2 0. 03 u zc 45 86 38 21 63 1 0. 11 17 .4 0 0. 94 77 .7 3 0. 10 47 .3 3 33 .7 8 99 .6 6 0. 52 8 0. 47 2 73 9 –1 .2 6 u zc 45 86 39 21 62 2 0. 13 14 .2 1 0. 60 80 .2 9 0. 18 0. 04 0. 06 0. 03 44 .0 3 40 .2 8 99 .5 5 0. 42 3 0. 57 7 67 8 –3 .2 8 u zc 45 86 40 21 60 2 0. 07 13 .4 5 0. 56 81 .7 8 0. 12 0. 01 0. 02 43 .6 7 42 .3 4 10 0. 24 0. 39 7 0. 60 3 62 0 –3 .2 7 u zc 45 86 42 21 55 2 0. 08 17 .1 3 0. 41 78 .1 4 0. 17 0. 01 0. 03 46 .7 6 34 .8 8 99 .4 7 0. 50 3 0. 49 7 72 6 –1 .3 4 u zc 45 86 43 21 50 4 0. 16 0. 13 7 16 .7 0 2. 32 9 0. 45 0. 10 3 80 .1 2 2. 25 1 0. 29 0. 08 2 0. 02 0. 00 0 0. 03 0. 02 0 0. 01 0. 00 4 0. 01 0. 00 0 46 .9 3 36 .8 7 10 1. 48 0. 48 3 0. 51 7 77 3 –0 .7 6 u zc 45 86 44 21 47 2 0. 75 5. 30 0. 39 88 .6 1 0. 15 0. 01 0. 02 37 .1 4 57 .1 9 10 0. 94 0. 15 8 0. 84 2 49 9 –0 .5 7 u zc 45 86 45 21 44 2 2. 21 5. 04 0. 48 87 .6 4 0. 09 0. 03 0. 03 0. 06 38 .8 5 54 .2 1 10 1. 00 0. 16 0 0. 84 0 46 6 –1 .0 9 u zc 45 86 46 21 44 2 0. 58 18 .6 3 0. 16 76 .4 1 0. 79 0. 02 0. 04 0. 02 0. 03 48 .2 0 31 .3 4 99 .7 9 0. 54 2 0. 45 8 81 6 –0 .8 0 u zc 45 86 57 20 30 3 0. 09 0. 05 3 19 .9 9 5. 86 7 0. 64 0. 25 1 74 .8 5 5. 58 0 0. 55 0. 15 4 0. 03 0. 00 1 0. 03 0. 01 9 0. 01 0. 00 4 0. 02 0. 00 2 48 .9 2 28 .8 1 99 .0 9 0. 59 4 0. 40 6 66 7 –2 .5 2 u zb 45 86 58 20 30 1 0. 05 19 .0 7 0. 57 76 .1 8 0. 35 0. 04 0. 01 0. 03 48 .2 9 30 .9 9 99 .4 0 0. 56 3 0. 43 7 63 3 –2 .8 1 u zb 87 .7 20 04 1 0. 10 28 .1 7 0. 57 66 .6 6 0. 45 0. 03 0. 01 0. 01 0. 03 56 .0 6 11 .7 7 97 .1 9 0. 84 5 0. 15 5 94 6 –1 .6 7 u zb 10 7. 7 19 84 2 0. 56 10 .8 3 2. 93 82 .0 3 0. 17 0. 02 0. 02 0. 02 42 .4 4 43 .9 8 10 0. 98 0. 37 2 0. 62 8 63 7 –1 .2 2 u zb 14 9. 8 19 42 3 0. 09 0. 06 7 27 .0 3 3. 44 8 1. 25 0. 38 9 69 .3 6 3. 57 7 0. 51 0. 00 9 0. 06 0. 00 0 0. 01 0. 00 9 0. 03 0. 00 3 0. 03 0. 04 6 55 .8 3 15 .0 3 99 .8 6 0. 81 7 0. 18 3 93 0 –1 .3 7 u zb 21 0. 9 18 81 2 0. 08 5. 49 1. 75 87 .0 9 0. 14 0. 03 0. 02 0. 02 36 .4 1 56 .3 1 10 0. 26 0. 17 5 0. 82 5 53 2 –0 .3 9 u zb 24 9. 2 18 43 4 0. 21 0. 29 2 10 .0 6 1. 39 0 2. 24 0. 27 2 83 .9 0 1. 36 1 0. 22 0. 02 0 0. 02 41 .2 5 47 .3 9 10 1. 38 0. 32 6 0. 67 4 69 3 –0 .2 1 u zb 29 0. 8 18 01 2 0. 03 8. 73 2. 09 85 .5 7 0. 21 0. 05 0. 02 39 .8 0 50 .8 5 10 1. 78 0. 27 8 0. 72 2 66 0 –0 .1 0 u zb 34 4. 1 17 48 3 0. 07 0. 01 3 13 .0 2 0. 65 6 2. 32 0. 22 3 80 .8 0 0. 41 1 0. 24 0. 02 8 0. 13 0. 03 1 0. 01 0. 00 8 0. 02 0. 03 43 .4 0 41 .5 5 10 0. 81 0. 42 1 0. 57 9 66 2 –1 .3 4 u zb 40 1. 3 16 91 2 0. 10 11 .9 1 2. 98 80 .2 7 0. 25 0. 25 0. 01 0. 01 0. 06 42 .0 9 42 .4 3 10 0. 08 0. 40 2 0. 59 8 61 7 –1 .8 2 u zb 46 1. 8 16 30 2 0. 07 15 .0 8 2. 98 75 .5 1 0. 31 0. 21 0. 01 0. 02 44 .3 0 34 .6 8 97 .6 5 0. 52 0 0. 48 0 71 4 –1 .3 4 u zb 49 1. 8 16 00 3 0. 07 0. 02 9 19 .5 8 1. 41 9 2. 88 0. 20 6 73 .8 6 1. 54 2 0. 32 0. 05 9 0. 38 0. 03 8 0. 02 0. 00 1 0. 02 0. 00 3 0. 03 0. 04 3 48 .8 6 27 .7 7 99 .9 3 0. 64 6 0. 35 4 73 8 –1 .7 0 u za 55 5. 5 15 36 4 0. 06 0. 01 9 12 .9 1 0. 59 9 3. 29 0. 46 3 77 .9 5 0. 53 3 0. 27 0. 04 5 0. 32 0. 13 5 0. 01 0. 00 4 0. 02 0. 00 7 42 .5 4 39 .3 5 98 .7 6 0. 44 8 0. 55 2 70 1 –1 .0 4 u za 58 5. 5 15 06 1 0. 03 12 .1 9 2. 90 79 .7 7 0. 23 0. 08 0. 02 42 .3 7 41 .5 6 99 .3 7 0. 41 5 0. 58 5 70 6 –0 .7 5 u za 64 8. 4 14 43 2 0. 06 12 .2 2 3. 35 80 .6 7 0. 24 0. 37 0. 03 0. 04 0. 07 42 .5 5 42 .3 5 10 1. 28 0. 41 3 0. 58 7 71 3 –0 .6 5 u za 68 9. 1 14 03 3 0. 09 0. 03 1 12 .9 6 0. 90 4 3. 30 0. 18 3 80 .1 3 0. 45 5 0. 28 0. 04 6 0. 38 0. 04 8 0. 02 0. 01 4 0. 05 0. 01 7 0. 04 0. 00 2 43 .2 8 40 .9 5 10 1. 36 0. 43 6 0. 56 4 72 5 –0 .6 8 u za 74 9. 2 13 43 1 0. 09 12 .3 0 2. 70 80 .3 9 0. 15 0. 48 0. 04 42 .3 0 42 .3 2 10 0. 37 0. 40 5 0. 59 5 72 3 –0 .5 1 u za 80 9. 2 12 82 3 0. 07 0. 03 2 11 .5 0 0. 58 3 3. 23 0. 35 5 80 .4 6 0. 95 2 0. 19 0. 05 4 0. 43 0. 12 2 0. 04 0. 02 5 0. 07 0. 01 3 0. 02 0. 00 2 41 .5 8 43 .2 1 10 0. 33 0. 39 1 0. 60 9 64 6 –1 .3 0 u za 85 4. 5 12 33 4 0. 05 0. 07 3 17 .0 4 2. 29 1 3. 35 0. 49 9 74 .1 2 0. 48 4 0. 31 0. 09 3 0. 48 0. 15 6 0. 07 0. 02 7 0. 04 46 .0 1 31 .2 3 98 .5 9 0. 58 7 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8 5. 31 10 0. 33 0. 05 1 0. 94 9 u za 94 1. 8 11 50 4 0. 01 0. 01 3 50 .4 8 0. 39 3 0. 07 0. 01 6 47 .4 5 0. 43 1 0. 57 0. 09 1. 17 0. 13 2 0. 02 0. 00 7 0. 02 0. 01 2 0. 02 0. 00 0 43 .4 2 5. 35 10 2. 77 0. 05 1 0. 94 9 m z 45 82 64 10 60 4 0. 01 0. 01 0 50 .6 6 0. 26 8 0. 07 0. 03 8 49 .1 1 0. 72 0 0. 48 0. 14 0. 92 0. 32 4 0. 01 0. 01 4 0. 03 0. 05 5 0. 05 0. 06 6 46 .3 4 6. 82 10 0. 73 0. 06 1 0. 93 9 m z 45 82 63 10 55 6 0. 04 0. 06 8 44 .6 2 14 .2 65 0. 70 1. 42 9 52 .4 8 11 .9 92 0. 38 0. 04 1. 73 0. 41 6 0. 01 0. 01 3 0. 03 0. 04 6 0. 05 0. 06 1 32 .9 2 15 .8 2 10 0. 19 0. 17 1 0. 82 9 m z 45 82 62 10 32 4 0. 21 0. 39 8 49 .1 7 1. 00 4 0. 12 0. 11 7 47 .1 6 0. 87 3 0. 44 0. 06 1. 44 0. 17 1 0. 01 0. 00 6 0. 01 0. 01 7 0. 05 0. 05 1 41 .7 2 6. 44 10 1. 19 0. 06 3 0. 93 7 m z 45 82 60 10 24 4 0. 02 0. 01 4 49 .1 2 0. 62 3 0. 12 0. 06 0 47 .5 2 1. 03 8 0. 37 0. 07 2. 40 0. 13 0 0. 01 0. 02 9 0. 03 0. 01 8 0. 04 0. 07 4 39 .6 4 8. 97 10 1. 66 0. 08 8 0. 91 2 m z 45 82 59 10 10 3 50 .8 0 0. 09 8 0. 07 0. 02 2 47 .7 1 0. 76 9 0. 39 0. 02 1. 77 0. 19 7 0. 01 0. 02 0 42 .1 7 6. 23 10 1. 47 0. 06 0 0. 94 0 m z 45 82 58 10 03 8 0. 01 0. 01 6 50 .8 8 0. 42 9 0. 10 0. 08 0 47 .7 7 0. 98 5 0. 46 0. 10 1. 57 0. 20 8 0. 01 0. 01 0 0. 02 0. 02 0 0. 02 0. 02 8 42 .2 4 5. 86 98 .7 9 0. 05 7 0. 94 3 (c on tin ue d) https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 38 of 139 geusbulletin.org ta bl e 9 (c on tin ue d) r ep re se nt at iv e co ex is tin g m ag ne tit e an d ilm en ite c om po si tio ns p er s am pl e ilm en ite zo ne   sa m pl e id st ra tig ra ph ic he ig ht n   si o 2 ti o 2 al 2o 3 fe o * m no m go ca o cr 2o 3 n io fe o   fe 2o 3   to ta l   x h e   x il       (m ) av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd m z 45 82 55 10 02 2 0. 02 51 .3 0 0. 05 47 .0 9 0. 52 2. 00 0. 01 0. 06 0. 09 41 .7 0 5. 66 10 0. 03 0. 05 5 0. 94 5 m z 45 82 54 98 8 1 49 .9 6 0. 08 46 .8 0 0. 53 2. 01 0. 01 0. 06 40 .7 5 6. 67 10 0. 50 0. 06 6 0. 93 4 m z 45 82 53 96 1 2 0. 01 50 .1 8 0. 27 46 .7 5 0. 44 2. 12 0. 06 41 .5 3 6. 61 10 2. 26 0. 06 4 0. 93 6 m z 45 82 51 95 3 5 0. 02 0. 02 6 49 .3 8 0. 49 2 0. 10 0. 04 9 49 .2 3 0. 38 1 0. 42 0. 06 2. 07 0. 13 6 0. 07 0. 02 7 0. 02 0. 03 6 38 .2 2 9. 41 10 2. 43 0. 09 6 0. 90 4 m z 45 82 50 93 5 7 0. 08 0. 18 1 51 .9 6 0. 64 7 0. 10 0. 05 7 46 .6 9 0. 43 3 0. 52 0. 11 2. 07 0. 08 7 0. 02 0. 02 0 0. 01 0. 01 5 0. 04 0. 07 3 42 .5 0 4. 60 10 1. 56 0. 04 4 0. 95 6 m z 45 82 49 93 0 8 0. 05 0. 09 0 51 .7 8 0. 55 2 0. 06 0. 02 8 46 .6 4 0. 73 4 0. 45 0. 10 2. 08 0. 20 4 0. 01 0. 02 0 0. 03 0. 04 3 41 .9 4 4. 62 10 1. 25 0. 04 5 0. 95 5 m z 45 82 47 92 0 8 0. 01 0. 01 1 52 .1 6 0. 53 1 0. 09 0. 09 5 46 .1 0 0. 66 5 0. 46 0. 04 1. 85 0. 15 3 0. 01 0. 01 5 0. 06 0. 06 2 0. 04 0. 06 8 44 .6 6 3. 45 98 .9 0 0. 03 2 0. 96 8 m z 45 82 89 89 0 2 0. 05 48 .1 5 0. 13 47 .7 6 0. 56 1. 59 0. 05 0. 24 0. 00 3 0. 04 40 .1 4 8. 84 10 2. 77 0. 08 7 0. 91 3 m z 45 82 88 87 7 3 0. 02 0. 02 5 50 .5 2 1. 00 5 0. 08 0. 00 6 48 .1 0 0. 35 8 0. 60 0. 08 2. 38 0. 12 0 0. 14 0. 01 1 0. 03 0. 02 2 41 .2 9 8. 52 10 2. 06 0. 08 1 0. 91 9 m z 45 82 87 87 5 3 0. 01 0. 00 3 49 .2 8 0. 96 5 0. 11 0. 03 5 48 .9 5 0. 40 2 0. 44 0. 08 2. 23 0. 10 0 0. 15 0. 00 4 0. 04 0. 07 3 37 .1 4 9. 42 10 0. 12 0. 09 8 0. 90 2 m z 45 82 86 84 7 2 0. 02 50 .9 8 0. 12 45 .6 2 0. 43 1. 97 0. 01 0. 01 0. 02 42 .2 5 4. 18 10 0. 31 0. 04 1 0. 95 9 lz c 45 82 85 83 0 3 0. 03 0. 01 1 50 .9 4 0. 84 9 0. 09 0. 01 7 46 .0 1 0. 60 3 0. 39 0. 07 2. 36 0. 04 1 0. 00 0. 00 5 0. 06 0. 09 4 42 .0 3 5. 47 10 2. 10 0. 05 3 0. 94 7 lz c 45 82 84 81 6 3 0. 15 0. 24 8 51 .3 6 0. 68 8 0. 16 0. 21 5 46 .9 5 1. 45 3 0. 47 0. 09 2. 34 0. 21 3 0. 06 0. 06 8 0. 02 0. 01 7 0. 04 0. 03 5 40 .2 9 5. 76 10 2. 31 0. 05 8 0. 94 2 lz c 45 82 83 80 8 3 0. 02 0. 01 5 52 .2 2 0. 19 9 0. 22 0. 18 3 45 .4 8 0. 77 8 0. 44 0. 03 3. 14 0. 16 1 0. 02 0. 02 6 0. 16 0. 02 8 0. 03 0. 03 2 42 .4 7 5. 30 97 .7 7 0. 05 0 0. 95 0 lz c 45 82 81 79 8 2 0. 05 47 .0 7 0. 14 47 .2 5 0. 45 2. 12 0. 04 0. 11 35 .8 9 9. 58 10 0. 16 0. 10 2 0. 89 8 lz c 45 82 82 75 2 3 0. 01 0. 01 2 50 .8 9 1. 03 7 0. 07 0. 00 2 44 .5 1 1. 49 7 0. 43 0. 05 3. 26 0. 20 0 0. 00 0. 00 0 0. 03 0. 06 0 41 .1 2 5. 81 10 2. 50 0. 05 6 0. 94 4 lz c 45 82 10 72 3 2 0. 08 51 .0 8 0. 09 48 .4 3 0. 54 2. 04 0. 16 0. 08 40 .3 9 7. 16 99 .0 3 0. 07 1 0. 92 9 lz c 45 82 78 71 0 7 0. 05 0. 05 2 48 .8 3 0. 31 2 0. 21 0. 30 6 48 .4 0 0. 46 9 0. 39 0. 06 2. 62 0. 11 0 0. 01 0. 01 6 0. 09 0. 03 5 0. 07 0. 05 4 39 .0 6 10 .7 0 10 1. 96 0. 10 4 0. 89 6 lz c 45 82 77 68 1 2 0. 20 49 .3 4 0. 51 48 .6 9 0. 47 1. 45 0. 05 0. 15 0. 02 40 .8 9 7. 93 10 1. 49 0. 07 8 0. 92 2 lz b 45 82 01 42 4 2 0. 01 47 .0 7 0. 11 49 .7 3 0. 45 2. 04 0. 20 36 .3 5 12 .1 5 10 1. 46 0. 12 5 0. 87 5 lz b 45 82 31 36 7 4 0. 02 0. 02 9 48 .0 2 0. 35 1 0. 16 0. 08 8 49 .8 4 0. 75 5 0. 38 0. 10 2. 49 0. 23 0 0. 02 0. 02 2 0. 27 0. 01 8 0. 06 0. 08 9 37 .2 1 12 .4 6 10 1. 44 0. 12 4 0. 87 6 lz b 45 82 26 29 4 3 0. 01 0. 02 2 48 .3 8 1. 51 8 0. 06 0. 05 0 48 .4 3 1. 54 1 0. 52 0. 13 2. 49 0. 18 0 0. 02 0. 02 0 0. 20 0. 00 3 0. 08 0. 07 0 36 .7 5 10 .6 0 10 0. 22 0. 10 9 0. 89 1 lz b 45 82 25 17 5 1 0. 04 48 .9 6 0. 14 46 .2 9 0. 28 3. 05 0. 05 0. 15 0. 19 38 .1 3 9. 11 97 .6 0 0. 09 1 0. 90 9 lz b 45 82 20 16 1 1 43 .1 1 0. 09 46 .6 7 0. 32 2. 33 0. 01 0. 27 0. 08 35 .7 7 14 .4 9 97 .6 0 0. 14 6 0. 85 4 lz a 45 82 18 12 5 2 48 .1 1 0. 11 47 .0 2 0. 46 2. 49 0. 09 0. 09 37 .5 0 9. 52 98 .2 0 0. 09 7 0. 90 3 lz a 45 82 15 10 7 4 0. 05 0. 04 4 45 .0 5 2. 12 0 0. 23 0. 20 6 47 .6 5 0. 99 3 0. 42 0. 10 2. 59 0. 08 0 0. 20 0. 04 0. 03 0. 02 2 34 .9 6 13 .2 9 97 .0 3 0. 13 8 0. 86 2 lz a 45 82 13 84 2 47 .6 8 0. 05 45 .2 5 0. 40 2. 86 0. 01 0. 06 0. 09 38 .9 6 9. 26 97 .4 7 0. 09 1 0. 90 9 lz a 45 82 12 54 2 0. 01 47 .2 1 0. 06 47 .2 9 0. 39 1. 40 0. 01 0. 14 0. 04 38 .9 0 8. 50 10 0. 14 0. 08 7 0. 91 3 n: n um be r of in di vi du al m in er al a na ly se s us ed fo r ea ch c al cu la te d av er ag e (a ve .) w ith 1 σ sd w he re n > 2 . f eo *, to ta l i ro n as f eo . i ro n is r ed is tr ib ut ed b et w ee n fe o a nd f e 2o 3 b as ed o n ch ar ge b al an ce (m ag ne tit e, 3 ca tio ns a nd 4 o xy ge ns ; ilm en ite 2 c at io ns a nd 3 o xy ge ns ). x il: m ol e fr ac tio n of il m en ite . x h e: m ol e fr ac tio n of h em at ite . x u lv : m ol e fr ac tio n of u lv ös pi ne l. x m t: m ol e fr ac tio n of m ag ne tit e. t em pe ra tu re (° c) a nd o xy ge n fu ca ci ty (l og fo 2 n n o ) f or c oe xi st in g ox id e m in er al s ar e ba se d on g hi or so & e va ns (2 00 8) . s tr at ig ra ph ic h ei gh t i s gi ve n as h ig h an d lo w r ef er rin g to th e in te rv al u se d fo r av er ag in g. d at a av ai la bl e in s up pl em en ta ry f ile s 1. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 39 of 139 geusbulletin.org et al. 1998; sonnenthal & mcbirney 1998). the slope of the linear variation in an content is 1.6 an% per 100 m of stratigraphy between an25 and an63 and confirms the findings of previous studies (fig. 1; wager & brown 1967; mcbirney 1989a). the plagioclase compositions at the appearance of ferrobustamite are similar (an32-33) to ubs plagioclase at a corresponding stratigraphic level (salmonsen & tegner 2013). the maximum an content of plagioclase in the hz (c. 65% an) reaches about the same value as found in the lza (maaløe 1976a; holness et al. 2015). there is a noticeably higher variability in the lza and lzb that significantly decreases in the mz and particularly in the uz (table 12), attributable to plagioclase zoning and an elevated residual melt content in the lower part of the intrusion (maaløe 1976a; tegner et al. 2009; humphreys 2009). the k content of plagioclase shows an irregular variation largely independent of the zone boundaries. the variation is irregular at around 0.015 pfu through lz to the base of the mz, then increases until the middle part of uzb, after which the content decreases into the uzc (fig. 26). there is an apparent reversal to lower values in the lower part of uza. the general increase in k in uz was also observed by jang & naslund (2001), using bulk analyses of separated plagioclase. the fe content on the other hand is within the range of 0.008 to 0.025 pfu in the lz and mz but increases from the base of uza, all the way into the uzc, despite some irregularities. the uz variation in fe of plagioclase was interpreted as a corresponding enrichment in the liquid by tegner (1997) and tegner & cawthorn (2010). 4.4.3 pyroxenes the clinopyroxene composition, expressed as mol% mg/ (mg+fe*) or mg# (magnesium number), shows a systematic upward increase in iron from mg# 71.6 ± 4.2 at the base of lza to mg# 0.8 ± 0.2 in uzc and at the sh (table 12; fig. 27). the mg# in clinopyroxene decreases upward, accelerating slightly in lzc and even more so in uzb, and continues into uzc. the variation compares well to predicted ideal fractional crystallisation trends in a closed basaltic system (maaløe 1976b; morse 1996). the slope of the trend in mg# varies from 1.3% per 100 m of stratigraphy in the lz and uza to 10% at the top of uzb. there are a few noticeable departures from the general upward increase in iron in the pyroxenes: at the base of uza, again at the base of uzb, and for the uzc where mg# rapidly levels off to zero. the reversal to higher mg# at the base of uzb was also observed for olivine (fig. 25). the titanium content varies widely up to 0.005 pfu, although with an initially weak increase through lza and lzb, followed by a decrease through lzc and the mz, and terminated by a systematic increase from uza to uzc (fig. 27). the aluminium content appears to be relatively constant in lza and lzb and then systematically decreases to uzc with few irregularities (except at the mz–uza boundary). variable ti and al results in lz–mz–uza uzc uzb cl f oh alvi + fe3+ fe2+ + mn mg cl al vi +f e 3+ oxides apatite mica mt ilm ulv he fe2+ fe3+ ti a b c fig. 24 compositions of minor mineral components of the ls. a: ternary fe2+–ti–fe3+ diagram showing the main molecular components of feti oxide solid solutions. the analysed oxides (blue dots) fall along two series approximately defining the ilmenite-hematite and ulvöspinel-magnetite solid solutions (grey dashed line). ilm: ilmenite. he: hematite. ulv: ulvöspinel. mt: magnetite. tie-lines are shown (grey solid lines) between average compositions (black dots) of the two series for individual thin sections. b: ternary diagram showing the monovalent cl–f–oh anion variation in apatite. solid blue dots: this study. open blue dots: previously published data (nash 1976; brown & peckett 1977; larsen 1992). apatite of the lz to uza is cl-bearing, the apatite of uzb is hydroxyl-bearing, and apatite in the uzc is f-bearing. c: ternary diagram showing the octahedral (alvi + fe3+) – mg – (fe2+ + mn) variation in mica (yellow dots). octahedral alumininum (alvi) is calculated assuming that tetrahedral al and fe3+ with si filling in the tetrahedral site to an ideal occupancy. the remaining iron is calculated as fe2+. biotite analysed in this study contains little, if any, alvi and fe3+. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 40 of 139 geusbulletin.org ta bl e 10 r ep re se nt at iv e ap at ite c om po si tio ns zo ne   st ra tig ra ph ic he ig ht n si o 2 (w t% ) fe o (w t% ) m no (w t% ) ca o (w t% ) p 2o 5 (w t% ) to ta l (w t% )   f (w t% ) cl (w t% ) at om ic m on ov al en t a ni on s f cl o h su m cl /f   (m ) av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd u bs 21 68 3 0. 33 0. 02 2 0. 34 0. 16 6 0. 01 0. 00 0 54 .9 2 0. 52 5 41 .6 7 0. 15 9 97 .2 7 3. 86 0. 20 2 0. 06 0. 03 0 1. 98 2 0. 08 7 0. 01 6 0. 00 8 0. 00 2 0. 00 6 2. 00 0 0. 00 8 u zc 21 65 1 0. 36 0. 30 0. 02 54 .8 9 40 .8 5 96 .4 1 3. 71 0. 15 1. 92 9 0. 04 3 0. 02 9 2. 00 0 0. 02 2 u zc 21 64 2 0. 53 0. 55 0. 03 54 .6 8 41 .4 2 97 .2 1 3. 75 0. 07 1. 92 7 0. 02 0 0. 05 3 2. 00 0 0. 01 0 u zc 21 63 3 0. 35 0. 00 1 0. 42 0. 01 9 0. 01 0. 00 0 54 .6 0 1. 33 1 41 .6 7 0. 67 9 97 .0 5 3. 71 0. 01 9 0. 08 0. 01 1 1. 91 2 0. 04 3 0. 02 1 0. 00 3 0. 06 7 0. 04 1 2. 00 0 0. 01 1 u zc 21 62 3 0. 30 0. 01 2 0. 27 0. 11 5 0. 06 0. 03 5 54 .5 8 0. 50 3 41 .7 2 0. 13 5 96 .9 2 3. 64 0. 06 0 0. 08 0. 01 5 1. 88 0 0. 03 3 0. 02 3 0. 00 4 0. 09 6 0. 03 4 2. 00 0 0. 01 2 u zc 21 60 6 0. 28 0. 02 2 0. 34 0. 07 1 0. 03 0. 02 4 54 .6 3 0. 32 5 41 .5 7 0. 28 8 96 .8 5 3. 75 0. 09 1 0. 10 0. 01 1 1. 93 4 0. 04 3 0. 02 6 0. 00 3 0. 04 0 0. 03 8 2. 00 0 0. 01 4 u zc 21 59 4 0. 37 0. 07 4 0. 47 0. 09 1 0. 05 0. 02 5 54 .1 8 0. 30 9 41 .8 4 0. 12 7 96 .8 9 3. 64 0. 06 0 0. 09 0. 01 9 1. 87 5 0. 03 1 0. 02 6 0. 00 5 0. 09 9 0. 03 1 2. 00 0 0. 01 4 u zc 21 55 3 0. 36 0. 28 5 0. 59 0. 16 0 0. 03 0. 00 2 54 .4 0 0. 47 6 41 .2 7 0. 39 3 96 .6 6 3. 78 0. 04 9 0. 04 0. 02 3 1. 95 6 0. 03 1 0. 01 2 0. 00 6 0. 03 2 0. 03 7 2. 00 0 0. 00 6 u zc 21 50 1 0. 30 0. 62 0. 03 55 .4 6 41 .9 8 98 .3 9 4. 08 0. 12 2. 06 3 0. 03 2 2. 09 5 0. 01 5 u zc 21 47 5 0. 23 0. 05 8 0. 37 0. 14 4 0. 04 0. 02 3 55 .2 2 0. 30 2 42 .1 4 0. 29 6 98 .0 0 3. 82 0. 08 5 0. 10 0. 02 9 1. 94 4 0. 03 9 0. 02 6 0. 00 8 0. 03 0 0. 04 8 2. 00 0 0. 01 4 u zc 21 44 4 0. 24 0. 15 2 0. 51 0. 10 8 0. 02 0. 02 5 55 .2 6 0. 59 2 41 .8 3 0. 32 5 97 .8 7 3. 88 0. 12 8 0. 13 0. 01 8 1. 97 9 0. 04 9 0. 03 5 0. 00 5 2. 01 4 0. 01 8 u zc 21 44 4 0. 28 0. 06 7 0. 31 0. 05 1 0. 03 0. 01 4 54 .5 4 0. 55 2 41 .7 2 0. 09 0 96 .8 7 3. 72 0. 04 8 0. 09 0. 02 4 1. 91 9 0. 02 0 0. 02 3 0. 00 7 0. 05 7 0. 02 4 2. 00 0 0. 01 2 u zc 21 41 5 0. 38 0. 20 0 0. 36 0. 09 7 0. 03 0. 01 8 55 .1 1 0. 79 8 41 .5 4 0. 39 2 97 .4 3 3. 75 0. 06 8 0. 16 0. 06 1 1. 92 4 0. 02 6 0. 04 5 0. 01 7 0. 03 0 0. 01 8 2. 00 0 0. 02 4 u zc 21 19 3 0. 11 0. 06 0 0. 41 0. 00 6 0. 02 0. 00 0 54 .6 0 0. 82 8 42 .2 3 0. 49 2 97 .3 7 3. 87 0. 06 8 0. 07 0. 00 0 1. 97 6 0. 01 9 0. 00 4 2. 00 0 0. 01 0 u zc 20 75 4 0. 13 0. 01 8 0. 32 0. 04 4 0. 05 0. 02 9 55 .6 2 0. 46 7 42 .5 7 0. 20 0 98 .6 9 3. 58 0. 21 3 0. 08 0. 02 0 1. 82 1 0. 10 3 0. 02 1 0. 00 5 0. 15 8 0. 10 7 2. 00 0 0. 01 1 u zc 20 60 3 0. 21 0. 07 4 0. 40 0. 06 1 0. 04 0. 00 7 54 .7 9 0. 40 8 41 .4 9 0. 28 2 96 .9 2 3. 73 0. 20 7 0. 12 0. 03 1 1. 92 8 0. 11 0 0. 03 2 0. 00 8 0. 04 0 0. 06 6 2. 00 0 0. 01 7 u zc 20 46 1 0. 12 0. 49 0. 10 56 .4 2 44 .9 5 10 2. 09 4. 01 0. 03 1. 94 9 0. 00 7 0. 04 3 2. 00 0 0. 00 4 u zb 20 32 3 0. 35 0. 15 9 0. 71 0. 28 5 0. 08 0. 02 3 55 .3 7 0. 21 1 41 .8 4 0. 27 3 98 .3 4 3. 68 0. 07 3 0. 05 0. 00 6 1. 88 1 0. 03 1 0. 01 4 0. 00 2 0. 10 5 0. 03 2 2. 00 0 0. 00 7 u zb 20 30 2 0. 28 0. 45 0. 05 54 .3 8 41 .6 9 96 .8 5 3. 61 0. 14 1. 86 7 0. 03 8 0. 09 5 2. 00 0 0. 02 0 u zb 20 04 3 0. 13 0. 04 2 0. 33 0. 07 8 0. 04 0. 03 9 54 .7 9 0. 56 0 41 .0 9 0. 79 4 96 .3 8 3. 91 0. 11 2 0. 09 0. 00 5 2. 02 5 0. 03 1 0. 02 6 0. 00 2 2. 05 1 0. 01 3 u zb 19 84 6 0. 31 0. 14 4 0. 61 0. 17 7 0. 06 0. 02 8 56 .6 8 0. 41 5 41 .4 0 0. 23 8 99 .0 6 1. 58 0. 35 0 0. 07 0. 00 9 0. 84 0 0. 18 0 0. 02 0 0. 00 2 1. 14 0 0. 18 1 2. 00 0 0. 02 4 u zb 19 42 3 0. 25 0. 09 3 0. 45 0. 05 3 0. 05 0. 01 2 56 .5 6 0. 14 2 41 .5 6 0. 10 3 98 .8 7 2. 14 0. 42 5 0. 11 0. 03 0 1. 12 6 0. 21 4 0. 03 2 0. 00 9 0. 84 3 0. 20 5 2. 00 0 0. 02 8 u zb 18 81 6 0. 28 0. 13 5 0. 34 0. 17 2 0. 05 0. 03 6 54 .8 1 0. 37 8 40 .7 5 0. 89 6 96 .2 3 3. 94 0. 12 0 0. 07 0. 01 2 2. 04 4 0. 07 5 0. 01 8 0. 00 3 2. 06 3 0. 00 9 u zb 18 43 5 0. 32 0. 13 5 0. 37 0. 17 2 0. 07 0. 00 0 57 .1 8 0. 85 0 41 .6 0 0. 78 6 99 .5 3 1. 90 0. 12 7 0. 07 0. 00 0 0. 99 7 0. 01 9 0. 98 4 2. 00 0 0. 01 9 u zb 18 01 6 0. 16 0. 04 1 0. 28 0. 10 5 0. 08 0. 04 7 55 .4 1 0. 85 5 44 .3 7 0. 79 6 10 0. 30 3. 90 0. 04 3 0. 07 0. 01 5 1. 92 9 0. 04 1 0. 01 7 0. 00 4 0. 05 4 0. 03 9 2. 00 0 0. 00 9 u zb 17 48 6 0. 18 0. 03 6 0. 51 0. 11 1 0. 06 0. 02 9 56 .5 7 0. 98 2 38 .2 7 0. 69 8 95 .6 0 1. 35 0. 24 8 0. 07 0. 02 1 0. 75 7 0. 12 9 0. 02 1 0. 00 7 1. 22 1 0. 12 7 2. 00 0 0. 02 8 u zb 16 91 8 0. 13 0. 07 6 0. 26 0. 21 1 0. 03 0. 00 0 55 .1 2 0. 88 2 41 .3 4 0. 55 8 96 .8 9 3. 96 0. 35 8 0. 06 0. 00 0 2. 03 7 0. 01 7 2. 05 4 0. 00 8 u zb 16 30 6 0. 14 0. 05 1 0. 29 0. 06 2 0. 05 0. 02 1 54 .4 6 0. 50 2 41 .3 8 0. 26 5 96 .3 3 3. 85 0. 08 8 0. 06 0. 00 8 1. 99 3 0. 04 5 0. 01 7 0. 00 2 2. 01 0 0. 00 9 lz b 63 4 1 0. 12 0. 33 0. 05 54 .7 3 42 .5 9 97 .8 2 2. 55 0. 41 1. 32 4 0. 11 5 0. 56 1 2. 00 0 0. 08 7 lz b 17 7 3 0. 14 0. 01 4 0. 27 0. 00 5 0. 05 0. 06 1 55 .0 8 0. 62 0 41 .6 9 0. 38 4 97 .2 3 2. 42 1. 40 5 0. 84 0. 06 6 1. 27 1 0. 09 7 0. 23 5 0. 02 1 0. 49 4 0. 11 8 2. 00 0 0. 18 5 lz b 17 5 1 0. 29 0. 57 0. 07 55 .0 7 40 .7 6 96 .7 6 2. 42 0. 75 1. 27 9 0. 21 2 0. 50 9 2. 00 0 0. 16 6 lz a 11 5 4 0. 17 0. 09 6 0. 20 0. 01 9 0. 09 0. 04 4 54 .6 7 0. 54 5 41 .2 3 0. 21 4 96 .3 6 2. 72 0. 13 5 0. 46 0. 00 9 1. 43 6 0. 07 9 0. 13 0 0. 00 3 0. 43 4 0. 08 2 2. 00 0 0. 09 0 n: n um be r of a na ly se s us ed fo r ea ch a ve ra ge (a ve .) w ith 1 σ sd fo r n > 2. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 41 of 139 geusbulletin.org a systematic increase in the ti/al ratio with a break in slope at the mz–uza boundary corresponding to the variations in the non-quadrilateral components (figs 19, 23). both total fe and mn increase throughout, while na shows some irregularities, but displays an overall up-section decrease, possibly interrupted by intervals showing constant or slightly increasing values (e.g. uza; fig. 27). in addition to the variation in the main quadrilateral components, the early studies of the skaergaard pyroxenes (muir 1951; brown et al. 1957; brown & vincent 1963) also observed strongly decreasing al and constant ti with fractionation for the clinopyroxene, corresponding to the variation observed in this study (figs 27, 28). nwe (1976) observed mg# 71–73 for clinopyroxene in the lower part of the cambridge drill core, which is within the  maximum range observed (table 5). clinopyroxene in  the mbs reaches mg# 74–79 (thy et al. 2008; humphreys & holness 2010; namur et al. 2013). the orthopyroxene composition matches that for olivine and clinopyroxene with a generally systematic upward increase in iron from mg# 71.6 ± 4.2 at the base of lza to 0.2 ± 0.3 at the top of the uzc and the sh (table 12). the inverted ferrobustamite shows limited variation reflecting their stratigraphically narrow occurrence in the uzc. 4.4.4 feti oxides both ilmenite and magnetite were analysed with a broad electron beam to obtain coexisting reconstituted mineral compositions approaching liquidus conditions. the success of this approach can be evaluated using the partitioning of mn and mg between the coexisting phases. bacon & hirschmann (1988) argued that the exchange coefficient calculated as kd  =  (mg/mn)mt  /  (mg/mn)il for liquidus conditions (il: ilmenite; mt: magnitite) were largely independent of temperature and composition and thus constant at 0.93. the kd (mg/mn) for this study varies widely at an average of 0.79 ± 0.40 with no systematic stratigraphic variations and encompasses the value for ideal liquidus conditions. the large standard deviation observed suggests limited success in obtaining liquidus oxide compositions, although it is possible that low-temperature, oxidation-exsolution of magnetite (buddington & lindsley 1964) may selectively affect mg and mn partitioning (bacon & hirschmann 1988). both ilmenite and magnetite show systematic upward decreases in magnesium from their appearances as primocrysts in lzc (fig. 29). despite large variations, other variables such as hematite in ilmenite show a systematic decrease from lzc into uza. ulvöspinel in magnetite also shows large variation but appears relatively constant until the upper part of uzb, where high oxidation can be observed. in the lza to lzb, where the feti oxides are late-crystallising phases, there is little ta bl e 11 in di vi du al a nd a ve ra ge m ic a co m po si tio ns zo ne sa m pl e id h ig ht (m ) n si o 2 ( w t% ) ti o 2 ( w t% ) al 2o 3 ( w t% ) fe o (w t% ) m no (w t% ) m go (w t% ) ca o (w t% ) n a 2o (w t% ) k 2o (w t% ) to ta l ( w t% ) at om ic r at io m g/ (f e+ m g) lz b 45 82 06 63 4 1 36 .9 3 5. 16 12 .7 3 20 .4 7 0. 22 12 .0 9 0. 21 0. 20 7. 98 95 .9 9 0. 51 3 lz b 45 82 21 17 5 3 35 .8 0 4. 59 13 .0 8 15 .6 9 0. 03 13 .8 5 0. 02 0. 15 8. 67 91 .8 8 0. 61 2 lz a 45 82 19 16 1 1 37 .1 8 4. 41 13 .7 4 11 .2 1 16 .7 5 0. 27 8. 73 92 .2 7 0. 72 7 lz a 45 82 15 11 5 1 36 .4 2 3. 64 14 .9 6 11 .5 6 0. 06 16 .8 3 0. 08 0. 27 8. 45 92 .2 6 0. 72 2 lz a 45 82 13 84 2 37 .1 4 4. 35 13 .8 0 14 .0 5 0. 05 15 .4 0 0. 36 8. 46 93 .6 2 0. 66 1 lz a 45 82 11 7 2 37 .4 6 4. 21 14 .8 1 12 .1 1 0. 02 17 .4 8 0. 06 0. 32 8. 23 94 .7 0 0. 72 0 in di vi du al a nd a ve ra ge s of m ic a w he n de te ct ed . n : n um be r of a na ly se s us ed fo r ca lc ul at in g av er ag es . https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 42 of 139 geusbulletin.org systematic compositional variation, except for a systematic increase for ulvöspinel for magnetite through lzb. 4.4.5 relative cryptic changes figures 25–29 show that the most dramatic changes in primocryst compositions are found in the uz. the relative changes in the composition of plagioclase, olivine and augite across the uz are shown in fig. 30 in terms of an content, fo content and pyroxene mg/(mg+fetotal) ratio. we refer to this as derivative cryptic variation that steadily increases up through uza and into uzb, reaching a maximum at 1962 m. above this, the changes reduce and eventually reach a constant value at the uzb–uzc boundary. all three silicates show parallel table 12 summary of cryptic primocryst compositions zone stratigraphic height (m) plagioclase olivine high-ca pyroxene low-ca pyroxene ave. sd min. max. ave. sd min. max. ave. sd min. max. ave. sd min. max. high low an an (1σ) an an fo fo (1σ) fo fo mg# mg# (1σ) mg# mg# mg# mg# (1σ) mg# mg# ubs 2169 2165 24.7 1.55 23.2 26.3 0.8 0.19 0.6 1.0 8.9 3.00 5.9 11.9 uzc 2165 2165 27.2 1.23 25.9 28.4 0.1 2.26 0.0 2.4 0.2 0.30 0.0 0.5 uzc 2165 2164 28.8 0.90 27.9 29.7 0.2 0.28 0.0 0.5 0.3 0.20 0.1 0.5 uzc 2163 2147 27.4 1.37 26.1 28.8 0.2 0.03 0.1 0.2 0.3 0.30 0.0 0.6 uzc 2144 2141 29.2 1.20 28.0 30.4 0.2 0.04 0.2 0.2 0.6 0.50 0.1 1.1 uzc 2091 2081 32.5 1.55 31.0 34.1 1.1 0.18 1.0 1.3 2.1 0.54 1.6 2.7 uzc 2075 2061 32.3 1.25 31.0 33.5 1.5 0.12 1.4 1.6 4.1 1.00 3.1 5.1 uzc 2060 2046 34.2 1.15 33.1 35.4 2.6 0.31 2.3 3.0 7.6 0.90 6.7 8.5 uzb 2044 2030 30.6 1.10 29.5 31.7 1.3 0.40 0.9 1.7 2.3 0.20 2.1 2.5 uzb 2024 1984 34.4 1.23 33.2 35.7 5.7 0.25 5.4 5.9 14.8 2.30 12.5 17.1 uzb 1962 1921 34.9 1.53 33.4 36.5 9.3 0.60 8.7 9.9 22.9 3.00 19.9 25.9 uzb 1902 1862 35.7 0.97 34.7 36.6 13.4 0.50 12.9 13.9 31.6 1.70 29.9 33.3 uzb 1843 1801 37.2 1.00 36.2 38.2 18.1 0.85 17.3 19.0 36.7 0.80 35.9 37.5 uzb 1782 1748 36.5 0.90 35.6 37.4 22.4 0.80 21.6 23.2 43.6 1.40 42.2 45.0 uzb 1728 1691 38.1 1.10 37.0 39.2 27.2 0.30 26.9 27.5 45.7 0.90 44.8 46.6 uzb 1671 1630 38.2 0.97 37.2 39.2 29.4 0.50 28.9 29.9 49.2 1.20 48.0 50.4 41.0 1.60 39.4 42.6 uzb 1620 1600 39.6 1.05 38.6 40.7 26.2 0.50 25.7 26.7 49.5 1.00 48.5 50.5 33.7 uza 1561 1527 40.7 0.95 41.7 39.8 28.7 0.35 28.3 29.0 49.7 1.50 48.2 51.2 uza 1506 1431 41.9 1.35 40.5 43.2 30.8 0.45 30.3 31.2 50.5 1.40 49.1 51.9 uza 1403 1343 43.4 1.35 42.1 44.8 34.7 0.50 34.2 35.2 53.9 2.30 51.6 56.2 43.8 uza 1323 1282 43.5 1.47 42.0 45.0 38.6 0.45 38.1 39.0 56.0 1.90 54.1 57.9 46.5 0.90 45.6 47.4 uza 1268 1237 43.9 2.20 41.7 46.1 37.8 0.60 37.2 38.4 55.7 0.00 55.7 55.7 uza 1218 1188 44.0 1.75 45.8 42.3 37.5 0.60 36.9 38.1 54.9 1.00 53.9 55.9 47.9 1.40 46.5 49.3 uza 1178 1131 44.3 1.40 42.9 45.7 39.4 0.60 38.8 40.0 56.5 3.60 52.9 60.1 47.7 0.50 47.2 48.2 mz 1098 1055 45.1 1.20 43.9 46.3 43.3 0.20 43.1 43.5 57.5 2.39 55.2 59.9 51.8 1.54 50.2 53.3 mz 1053 1032 46.4 1.65 44.7 48.0 60.2 2.44 57.8 62.6 54.9 2.31 52.6 57.2 mz 1024 1009 47.0 1.10 45.9 48.1 44.2 0.60 43.6 44.8 59.7 3.56 56.2 63.3 54.1 1.34 52.7 55.4 mz 1003 1002 47.6 1.73 45.9 49.4 58.7 2.56 56.1 61.3 52.8 2.54 50.2 55.3 mz 988 961 48.3 1.60 46.7 49.9 48.6 1.00 47.6 49.6 59.0 3.29 55.7 62.3 54.9 1.64 53.3 56.5 mz 953 930 48.2 1.80 46.4 50.0 61.8 1.96 59.9 63.8 55.0 1.11 53.9 56.1 mz 921 900 51.9 3.63 48.3 55.6 45.9 0.60 45.3 46.5 60.7 1.64 59.1 62.3 55.3 mz 890 875 50.0 1.60 48.4 51.6 50.5 0.70 49.8 51.2 62.8 1.77 61.1 64.6 lzc 847 816 50.5 1.70 48.8 52.2 53.9 1.70 52.2 55.6 63.6 2.91 60.7 66.5 59.8 lzc 808 798 49.9 1.53 48.3 51.4 55.1 0.60 54.5 55.7 64.4 1.78 62.6 66.2 58.4 lzc 784 742 50.9 1.67 49.2 52.5 65.3 3.86 61.5 69.2 60.7 0.62 60.1 61.4 lzc 723 703 53.2 1.37 51.9 54.6 52.2 0.45 51.8 52.7 65.1 1.84 63.3 66.9 58.7 lzb 703 681 51.9 1.50 50.4 53.4 51.5 0.45 51.1 52.0 64.6 2.06 62.5 66.6 55.5 lzb 634 580 54.8 1.73 53.1 56.5 49.8 0.40 49.4 50.2 65.4 3.30 62.1 68.7 lzb 558 488 55.3 2.77 52.6 58.1 52.4 0.50 51.9 52.9 65.2 2.25 63.0 67.5 61.7 lzb 447 367 56.4 2.37 54.0 58.8 53.4 0.50 52.9 53.9 65.5 3.33 62.1 68.8 58.0 2.02 56.0 60.0 lzb 346 221 58.9 1.70 57.2 60.6 57.0 0.70 56.3 57.7 68.4 3.78 64.6 72.2 62.9 lzb 177 173 60.2 1.53 58.6 61.7 53.0 0.90 52.1 53.9 68.0 5.23 62.8 73.2 60.8 lza 161 137 58.1 3.97 54.1 62.0 56.6 0.60 56.0 57.2 70.4 4.59 65.8 75.0 64.9 0.98 63.9 65.9 lza 125 96 60.3 3.33 56.9 63.6 58.0 0.95 57.1 59.0 70.9 0.91 70.0 71.8 66.7 0.37 66.3 67.0 lza 84 27 62.0 5.43 56.5 67.4 57.4 2.15 55.2 59.5 71.0 6.20 64.8 77.2 68.4 lza 7 63.2 5.65 57.5 68.8 61.1 3.60 57.5 64.7 71.6 4.20 67.4 75.8 64.4 calculated as averages (ave.) of three consecutive samples with 1σ level sd calculated for n < 2. min.: minimum. max.: maximum. stratigraphic height is given as high and low referring to the interval used for averaging. complete data set in supplementary file s1. an and fo in mol%. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 43 of 139 geusbulletin.org variation with the lowest values for plagioclase, then olivine and highest for clinopyroxene. a slight increase resumes from the base of the uzc, reaching a maximum in the centre of the uzc, and subsequently decreases until the sh. the uzc variation resembles the symmetric pattern of a sill-like segregation cap cooled both from the base and from the top. 4.4.6 co-variation of silicate minerals the composition of mafic silicate minerals is shown in fig. 31 as a function of coexisting plagioclase primocrysts. for similar an contents, the mg/(mg+fetotal) ratio is highest for clinopyroxene, followed by orthopyroxene and finally lowest for olivine. mafic minerals are shown in fig. 32 as a function of clinopyroxene composition, where forsterite of olivine defines a polynomial curve, between zero and c. 70% mg#. orthopyroxene with higher mg# shows an apparent linear relationship only reaching values of just below 50% mg#. 4.5 bulk-rock compositions the composition of the gabbros is principally controlled by modal proportions of primocrysts formed onto the solid substrata or walls of the chamber. although the individual phase compositions may reflect the coexisting melt composition, or equilibrium, their relative proportions reflect selective accumulation and redistribution processes operating during the early stages of gabbro formation. the spectacular and complex layering preserved throughout the intrusion reflects such processes, although their detailed explanations have been debated (wager & brown 1967; irvine 1979; maaløe 1978; hunter 1987; namur et al. 2015). it is likely that the melt component of the crystal mush, which was originally formed and trapped on the magma floor, may have been either retained or migrated upwards into the overlaying mush. here it may have crystallised, altering interstitial melt composition and possibly reacting with the solid components, 100 300 500 700 900 1100 1300 1500 1700 1900 2100 0 20 40 60 80 0.02 0.04 0.06 0.08 == olivine (fo mol%) mn (pfu) sh/ubs 90 -2 2 uzc uzb uza mz lzc lzb lza 0 st ra ti gr ap hi c he ig ht (m ) olivine corona structure all analyses ave. & sd fig. 25 cryptic variation of olivine as a function of stratigraphic height (m) in the ls. shown are forsterite content (fo mol%) and the numbers of mn pfu (4 oxygen). small blue dots: individually analysed points. solid black dots: average of each thin section with 1s sd. open blue dots: composition of coronas. red vertical dashed line: location of drill core 90-22. blue horizontal dashed lines: divisions of the ls. abbreviations in fig. 2. ave.: average. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 44 of 139 geusbulletin.org st ra ti gr ap hi c he ig ht (m ) plagioclase (an mol%) fetotal (pfu) sh/ubs uzc uzb uza mz lzc lzb lza plagioclase 90-22 k (pfu) 0 0.01 0.02 0.03 0.04 0.05 100 300 500 700 900 1100 1300 1500 1700 1900 2100 0 0.2 0.4 0.6 0.8 0 0.01 0.02 0.03 0.04 ave. & sd all analyses fig. 26 cryptic variation in plagioclase as a function of stratigraphic height (m) in the ls. shown are anorthite content (an mol%) and the numbers of k and fetotal cations pfu (8 oxygen). iron is given as total fe. blue horizontal dashed lines: divisions of the ls. symbols and abbreviations in figs 2, 25. 100 300 500 700 900 1100 1300 1500 1700 1900 2100 0.0 0.2 0.4 0.6 0.8 ti 0.02 0.04 0.06 0.04 0.08 0.12 0.50 1.00 1.50 ti/al atomic ratio augite st ra ti gr ap hi c he ig ht (m ) ti (pfu)mg/(mg + fe) ratio al (pfu) sh/ubs uzc uzb uza mz lzc lzb lza 90-22 ferrobustamite (ave. & sd) ave. & sd all analyses 0.00 0.00 0.00 fig. 27 cryptic variation in pyroxenes (augite) as a function of stratigraphic height (m) in the ls (continued in fig. 28). shown are the atomic ratio mg/(mg + fe), the numbers of ti and al cations pfu (4 cation and 6 oxygen) and the ti/al atomic ratio. iron is given as total fe. green open squares: ferrobustamite. blue horizontal dashed lines: divisions of the ls. other symbols and abbreviations in figs 2, 25. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 45 of 139 geusbulletin.org and thus complicating the understanding of gabbro compositions. despite the possibility for such late-stage modification, the overall composition of the gabbros will mainly record the proportions of primocryst in addition to a lesser amount of trapped melt mostly crystallised as the same minerals and in the same proportions. major elements are in the following sections defined as elements that are essential for forming the main primocrysts, while trace elements are defined as elements not essential to building the structure of the primocrysts, with some exceptions as we show in the following sections. the trace elements are discussed as two different groups, either included in the silicate building blocks or excluded from them and residing in the melt. 4.5.1 major elements this section will only describe the stratigraphic changes in the gabbro compositions that can be related to the principal changes in primocryst assemblages for the individual zone and subzone divisions. the major element composition of the gabbros is illustrated in fig. 33 and summarised in table 13. the average zone compositions for the composite profile are shown in table 14 and compare remarkably well to similar lateral intrusion-wide estimates by mcbirney (1989a). the volatile component expressed as loi was only determined for the surface samples (n = 86) for which the average is 0.64 ± 0.14% in the lz and mz and rising to 1.22 ± 0.38% in the uzc (not shown). the fe2o3/feo ratio (also not shown) is relatively constant throughout lza and lzb at 0.17 ± 0.07, then increases markedly in lzc to 0.35 ± 0.10, after which the ratio steadily decreases to 0.12 ± 0.02 in the lower part of uzc, and increases to 0.33 ± 0.12 in the upper part of the uzc and sh. this variation closely mimics the appearance of magnetite and its modal decrease upward and suggests only a modest secondary oxidation for the majority of the profile. the exception is the upper part of the uzc, which may indicate some late-stage or secondary oxidation. sh/ubs uzc uzb uza mz lzb lza 90-22 0.01 0.02 0.03 0.04 0.05 0.01 0.02 0.03 0.040.25 0.50 0.75 1.00 1.25 1.50 100 300 500 700 900 1100 1300 1500 1700 1900 2100 augite st ra ti gr ap hi c he ig ht (m ) fe (pfu) mn (pfu) na (pfu) lzc 0.00 0.000.00 ferrobustamite (ave. & sd) ave. & sd all analyses fig. 28 cryptic variation in pyroxenes (augite) as a function of stratigraphic height (m) in the ls (continued from fig. 27). shown are the numbers of fe, mn and na cations pfu (4 cation and 6 oxygen). iron is given as total fe. green open squares: ferrobustamite. blue horizontal dashed lines: divisions of the ls. other symbols and abbreviations in figs 2, 25. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 46 of 139 geusbulletin.org the variation of many of the major elements clearly shows the onset of feti oxides as a primocryst at the lzb–lzc boundary (fig. 33). this is as expected, particularly where tio2 abruptly increases, fe2o3, feo increases and sio2 decreases. the modal composition suggests that ilmenite is the dominating feti oxide, and that magnetite occurs in lower proportions than ilmenite (fig. 15). the transition from lza to lzb is weakly defined by a lower cao content in the former, which corresponds to the observed lower modal content of clinopyroxene in lza (fig. 15) and supports the traditional definition of the lza as lacking clinopyroxene as a primocryst (wager & brown 1967). 0 200 400 600 800 1000 1200 1400 1600 1800 2000 2200 0 0.05 0.10 0.15 0.05 0.10 0.15 mg (pfu) 0.05 0.10 0.15 0.20 magnetiteilmenite mg (pfu) ilmenite hematite fraction 0.2 0.4 0.6 0.8 1.0 ulvöspinel fraction magnetite st ra ti gr ap hi c he ig ht (m ) sh/ubs 90-22 uzc uzb uza mz lzc lzb lza 0 0 0 fig. 29 cryptic variation in ilmenite (blue) and magnetite (black) as a function of stratigraphic height (m) in the ls. shown are the numbers of mg cations pfu (ilmenite calculated as 2 cations and 3 oxygens and magnetite as 3 cations and 4 oxygens) as individual point analyses. shown are numbers of mg cations pfu in ilmenite and magnetite, mole fractions of hematite in ilmenite, and mole fraction of ulvöspinel in magnetite. compare to fig. 24 and table 9. red vertical dashed line: location of drill core 90-22. blue horizontal dashed lines: divisions of the ls. abbreviations in figs 2, 25. 1600 1700 1800 1900 2000 2100 2200 –0.01 0 0.01 0.02 0.03 0.04 0.05 sh/ubs uzb uzc uza plagioclase augite olivine st ra ti gr ap hi c he ig ht (m ) derivative cryptic variation olivine plagioclase augite fig. 30 derivative cryptic variation in the uzb and uzc for coexisting olivine (as fo mol%), plagioclase (as an mol%) and augite (as mg/(mg+fetotal) atomic ratio). averages are calculated as fractions in steps of four consecutive determinations. 0 20 40 60 80 20 30 40 50 60 70 olivine clinopyroxene orthopyroxene m g/ (m g + fe to ta l ) ( % ) an (mol.%) of plagioclase fig. 31 mg/(mg + fetotal) content in coexisting mafic minerals (olivine, clinopyroxene, and orthopyroxenes; %) vs. an (mol%) of coexisting plagioclase. the horizontal and vertical bars are the 1σ sd of repeated analyses in individual thin sections. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 47 of 139 geusbulletin.org the appearance of apatite as a primocryst at the uza–uzb boundary is likewise strongly reflected in the p2o5 content, which is low below the boundary, often just above the detection limit, but abruptly increases to nearly 4 wt% at the base of uzb. concurrent with the increase in p2o5 is a smaller rise in cao that can also be understood in terms of apatite accumulation. manganese shows a weak rise through uzb and followed by a marked rise in uzc that can be correlated to the appearance of inverted ferrobustamite in this subzone (fig. 20). the gabbro fluorine content is closely positively correlated with the modal amount of apatite (see supplementary file s1). the transition from uza to uzb, where sio2 and al2o3 drop and feo and cao rise, cannot directly be related to the appearance of primocryst apatite, even at the maximum amount of 10%. they can, however, be related to the high amounts of both iron-rich olivine and apatite observed at the base of the uzb as shown by the olivine and apatite modes (fig. 14; see also sections 5.7.7 and 5.11.8 on the effects of liquid immiscibility). the mz also shows large variations, but these cannot univocally be related to the reaction of olivine with interstitial melt to form orthopyroxene 0 10 20 30 40 50 60 70 80 0 10 20 30 40 50 60 70 80 y = 1.16x – 16.0 (r2 = 0.921) mg/(mg + fetotal) of cpx (%) m g/ (m g + fe to ta l ) o f o l, o px (% ) y = 0.555 + 0.081x + 0.011x2 (r2 = 0.988) 1:1 orthopyroxene olivine fig. 32 mg/(mg + fetotal) content in coexisting olivine (ol) and orthopyroxene (opx; %) vs. mg/(mg + fetotal) content in coexisting clinopyroxenes (cpx; %). equations for the best fits to olivine and orthopyroxene are given. 200 400 600 800 1000 1200 1400 1600 1800 2000 2200 20 30 40 50 60 sio2 tio2 al2o3 5 10 15 200 400 600 800 1000 1200 1400 1600 1800 2000 2200 5 10 15 20 25 30 fe2o3 feo 200 400 600 800 1000 1200 1400 1600 1800 2000 2200 mgo 0.2 0.4 0.6 0.8 1.0 mno cao na2o k2o 200 400 600 800 1000 1200 1400 1600 1800 2000 2200 p2o5 oxides (wt%) st ra ti gr ap hi c he ig ht (m ) lza lzb lzc mz uza uzb uzc sh/ubs lza lzb lzc mz uza uzb uzc sh/ubs 90-22 0 0 0 0 5 10 15 20 255 10 15 20 25 1 2 3 40.2 0.4 0.6 0.8 1.0 1.22 4 6 8 10 12 14 160 0 0 0 000 0 1 2 3 4 5 fig. 33 major elements of the ls gabbros, shown as oxides (wt%) vs. stratigraphic height (m). blue horizontal dashed lines: divisions of the ls. abbreviations in fig. 2. red vertical dashed line: location of drill core 90-22. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 48 of 139 geusbulletin.org ta bl e 13 s um m ar y of g ab br o m aj or e le m en t c om po si tio ns zo ne st ra tig ra ph ic he ig ht (m ) f si o 2 (w t% ) ti o 2 (w t% ) al 2o 3 (w t% ) fe 2o 3 (w t% ) fe o (w t% ) m no (w t% ) m go (w t% ) ca o (w t% ) n a 2o (w t% ) k 2o (w t% ) p 2o 5 (w t% ) to ta l (w t% ) hi gh lo w av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd u bs 21 69 21 65 51 .8 2 2. 38 0 1. 97 0. 53 0 10 .4 3 1. 16 1 5. 48 2. 27 2 15 .2 8 3. 29 8 0. 26 0. 08 0 0. 72 0. 63 2 7. 08 0. 77 0 3. 61 0. 48 6 0. 98 0. 09 9 0. 61 0. 31 8 98 .2 3 u zc 21 65 21 65 0. 00 0 50 .1 7 3. 47 7 1. 87 0. 49 2 9. 40 2. 39 7 6. 28 0. 84 4 18 .4 5 4. 75 9 0. 19 0. 01 1 0. 03 0. 01 6 7. 89 1. 06 3 3. 25 0. 94 8 0. 80 0. 15 6 0. 36 0. 20 2 98 .7 0 u zc 21 65 21 64 0. 00 0 47 .3 7 1. 45 3 2. 69 0. 91 8 7. 28 0. 64 2 7. 02 1. 09 4 20 .8 8 1. 86 9 0. 25 0. 02 6 0. 02 0. 00 5 8. 71 0. 64 5 2. 52 0. 21 0 0. 77 0. 09 5 0. 67 0. 14 0 98 .1 9 u zc 21 63 21 47 0. 00 1 48 .1 1 2. 42 2 2. 46 0. 50 5 7. 83 1. 72 9 6. 21 1. 92 4 20 .6 1 1. 17 2 0. 32 0. 12 9 0. 03 0. 02 3 8. 55 0. 45 3 2. 69 0. 61 2 0. 68 0. 10 0 0. 63 0. 21 4 98 .1 1 u zc 21 44 21 41 0. 00 3 45 .7 7 2. 24 1 2. 57 0. 15 3 6. 64 0. 87 1 5. 18 1. 50 3 24 .8 2 3. 12 0 0. 48 0. 07 8 0. 07 0. 01 5 9. 36 0. 90 6 2. 24 0. 45 3 0. 56 0. 19 0 0. 75 0. 10 5 98 .4 4 u zc 20 91 20 81 0. 01 0 45 .1 5 0. 65 2 2. 70 0. 24 4 8. 32 0. 42 9 2. 88 0. 32 3 24 .7 3 0. 65 5 0. 72 0. 01 3 0. 26 0. 04 0 10 .3 8 0. 31 6 2. 52 0. 13 0 0. 26 0. 03 3 0. 78 0. 04 0 98 .7 1 u zc 20 75 20 60 0. 01 2 44 .5 1 0. 48 4 2. 78 0. 17 3 7. 43 0. 38 7 3. 09 0. 64 7 26 .1 3 1. 54 0 0. 84 0. 05 4 0. 36 0. 08 4 10 .2 6 0. 06 0 2. 32 0. 13 8 0. 27 0. 03 7 0. 87 0. 07 1 98 .8 5 u zc 20 60 20 46 0. 01 4 45 .6 4 0. 91 4 2. 54 0. 26 2 8. 98 1. 39 1 2. 68 0. 31 1 23 .9 6 1. 85 2 0. 74 0. 24 4 0. 68 0. 33 2 9. 65 0. 42 7 2. 68 0. 35 9 0. 32 0. 03 3 0. 71 0. 06 7 98 .5 7 u zb 20 44 20 30 0. 01 6 51 .6 9 8. 59 6 2. 02 0. 61 8 9. 15 0. 20 0 4. 77 1. 65 1 18 .2 3 7. 24 1 0. 49 0. 19 0 0. 39 0. 40 2 7. 31 2. 85 8 3. 35 0. 77 4 0. 53 0. 24 7 0. 53 0. 21 6 98 .4 6 u zb 20 24 19 84 0. 01 9 44 .7 5 1. 78 3 2. 83 0. 37 2 10 .0 0 0. 32 2 4. 17 0. 73 3 22 .2 1 1. 14 7 0. 46 0. 01 7 1. 37 0. 27 7 9. 40 0. 27 9 2. 89 0. 13 9 0. 31 0. 04 6 0. 90 0. 29 6 99 .2 9 u zb 19 62 19 21 0. 02 9 44 .0 6 0. 73 7 3. 01 0. 12 9 10 .8 3 0. 34 2 4. 27 0. 85 4 20 .8 8 0. 42 2 0. 42 0. 00 7 1. 95 0. 23 4 9. 13 0. 35 9 3. 01 0. 12 2 0. 32 0. 03 3 1. 18 0. 06 4 99 .0 5 u zb 19 02 18 62 0. 03 9 44 .1 0 0. 66 4 3. 29 0. 11 0 12 .1 4 0. 25 4 3. 73 0. 43 5 19 .1 1 0. 79 0 0. 35 0. 00 7 2. 35 0. 19 2 8. 88 0. 06 8 3. 30 0. 09 3 0. 35 0. 04 0 1. 46 0. 09 3 99 .0 6 u zb 18 43 18 01 0. 05 0 41 .6 9 1. 85 5 3. 91 0. 34 3 10 .8 2 1. 50 1 3. 71 0. 74 3 20 .9 0 2. 23 2 0. 35 0. 03 9 3. 46 0. 61 6 9. 18 0. 05 1 2. 89 0. 43 7 0. 27 0. 06 1 1. 87 0. 15 8 99 .0 4 u zb 17 82 17 48 0. 06 2 41 .2 6 0. 61 3 4. 15 0. 11 0 10 .3 8 0. 06 6 2. 79 0. 10 9 21 .4 0 0. 38 6 0. 33 0. 00 7 4. 14 0. 27 1 9. 40 0. 21 3 2. 75 0. 01 6 0. 35 0. 11 7 2. 06 0. 30 1 99 .0 2 u zb 17 28 16 91 0. 07 3 37 .7 3 1. 92 6 5. 24 1. 05 8 10 .1 3 0. 14 2 3. 37 0. 37 9 22 .2 1 0. 67 0 0. 33 0. 01 1 4. 53 0. 40 5 9. 56 0. 60 9 2. 61 0. 02 2 0. 21 0. 02 1 3. 12 0. 83 2 99 .0 3 u zb 16 71 16 30 0. 08 6 40 .7 7 2. 84 4 4. 34 1. 16 5 10 .8 0 1. 17 8 3. 19 0. 94 7 18 .4 1 2. 36 9 0. 29 0. 02 8 4. 86 0. 59 0 10 .6 0 0. 63 6 2. 79 0. 34 4 0. 23 0. 04 4 2. 84 0. 06 1 99 .1 3 u zb 16 20 16 00 0. 09 7 42 .3 6 3. 15 3 4. 79 0. 33 5 10 .7 4 1. 15 8 4. 70 0. 75 5 16 .9 7 2. 00 9 0. 27 0. 02 7 4. 80 0. 46 4 10 .3 6 1. 04 3 2. 66 0. 30 2 0. 22 0. 05 2 1. 27 1. 93 9 99 .1 3 u za 15 61 15 27 0. 11 2 45 .9 3 0. 53 8 4. 66 0. 37 4 13 .2 8 0. 48 8 3. 20 1. 26 4 15 .0 6 1. 49 9 0. 25 0. 02 1 4. 43 0. 39 9 8. 93 0. 40 0 3. 21 0. 12 9 0. 24 0. 04 2 0. 06 0. 01 7 99 .2 5 u za 15 06 14 31 0. 12 6 45 .3 0 1. 77 8 5. 28 0. 90 4 12 .5 5 1. 29 4 3. 52 0. 88 0 14 .5 4 2. 71 4 0. 26 0. 03 2 5. 09 1. 10 1 9. 41 0. 46 5 2. 96 0. 38 5 0. 25 0. 11 9 0. 04 0. 00 9 99 .2 0 u za 14 03 13 43 0. 15 4 45 .3 8 3. 09 9 4. 94 2. 22 0 13 .0 4 1. 15 8 3. 21 0. 98 7 14 .5 5 1. 68 1 0. 24 0. 01 9 5. 32 0. 18 6 9. 50 0. 55 7 2. 93 0. 23 8 0. 20 0. 01 4 0. 05 0. 00 0 99 .3 8 u za 13 23 12 82 0. 17 7 43 .9 3 0. 51 9 5. 55 0. 24 9 11 .9 7 0. 11 3 3. 85 0. 03 7 15 .5 3 0. 55 8 0. 26 0. 00 0 6. 12 0. 13 4 9. 36 0. 28 7 2. 63 0. 02 6 0. 16 0. 00 0 0. 05 0. 00 0 99 .4 1 u za 12 68 12 37 0. 19 4 44 .2 8 4. 50 7 5. 11 1. 99 9 11 .9 9 0. 95 5 4. 65 1. 74 5 14 .4 6 2. 97 3 0. 25 0. 01 7 5. 81 0. 21 5 9. 89 1. 00 7 2. 60 0. 27 7 0. 17 0. 04 9 0. 06 0. 04 4 99 .2 9 u za 12 18 11 88 0. 21 0 45 .5 2 1. 46 3 5. 58 0. 43 9 12 .0 1 0. 56 3 3. 44 1. 14 2 13 .7 8 0. 84 0 0. 25 0. 00 9 5. 95 0. 30 0 10 .1 2 0. 66 8 2. 64 0. 13 5 0. 20 0. 01 4 0. 08 0. 00 0 99 .5 5 u za 11 78 11 31 0. 22 4 44 .1 4 0. 64 9 5. 79 0. 58 9 11 .9 1 0. 63 2 4. 88 1. 46 8 13 .7 9 0. 48 5 0. 25 0. 00 9 5. 96 0. 25 0 9. 79 0. 31 4 2. 53 0. 11 5 0. 18 0. 01 4 0. 06 0. 01 1 99 .2 7 m z 10 98 10 55 0. 25 1 42 .5 3 3. 03 6 7. 86 1. 58 8 13 .5 4 3. 93 2 4. 32 1. 10 4 13 .8 1 3. 62 1 0. 22 0. 06 0 4. 88 1. 80 2 9. 16 0. 16 6 2. 80 0. 85 8 0. 17 0. 05 0 0. 05 0. 01 3 99 .3 5 m z 10 53 10 32 0. 26 7 41 .6 5 4. 13 5 9. 97 3. 82 8 9. 76 0. 93 4 2. 65 3. 95 5 15 .6 0 1. 71 8 0. 28 0. 00 7 7. 20 0. 14 9 10 .2 1 0. 83 6 1. 96 0. 22 0 0. 11 0. 02 9 0. 03 0. 01 6 99 .4 1 m z 10 24 10 09 0. 27 7 43 .3 1 3. 19 4 6. 11 2. 33 0 11 .1 2 0. 14 9 3. 98 0. 92 2 15 .0 0 1. 35 2 0. 26 0. 00 0 7. 09 0. 71 4 10 .2 0 0. 59 2 2. 18 0. 05 0 0. 13 0. 01 8 0. 04 0. 01 4 99 .4 3 m z 10 03 10 02 0. 28 5 41 .1 0 6. 23 5 7. 91 3. 67 1 12 .6 9 3. 99 2 4. 88 2. 17 5 14 .8 3 4. 68 1 0. 23 0. 06 1 5. 33 1. 46 5 9. 78 1. 15 1 2. 45 0. 85 0 0. 14 0. 05 3 0. 04 0. 01 6 99 .3 8 m z 98 8 96 1 0. 29 1 39 .1 3 4. 19 0 9. 87 3. 10 0 8. 45 2. 07 8 5. 21 1. 48 9 17 .0 1 1. 54 1 0. 29 0. 02 7 7. 72 1. 59 3 10 .0 2 1. 79 9 1. 58 0. 40 4 0. 10 0. 01 7 0. 03 0. 00 5 99 .4 2 m z 95 3 93 0 0. 30 4 38 .1 2 4. 57 3 9. 71 1. 93 5 11 .4 3 1. 43 6 5. 88 2. 11 1 16 .3 0 3. 13 8 0. 24 0. 02 2 5. 75 0. 59 9 9. 47 1. 07 2 2. 10 0. 34 4 0. 12 0. 02 8 0. 03 0. 01 5 99 .1 4 m z 92 1 90 0 0. 31 6 44 .1 0 5. 15 5 5. 89 3. 72 8 14 .5 8 3. 57 9 2. 96 2. 14 0 12 .4 8 3. 98 0 0. 19 0. 04 6 5. 73 0. 47 7 10 .7 7 1. 06 4 2. 55 0. 44 6 0. 13 0. 03 9 0. 02 0. 01 3 99 .4 1 m z 89 0 87 5 0. 32 9 39 .6 3 2. 47 2 9. 05 1. 34 6 12 .8 0 1. 36 1 5. 62 1. 02 3 14 .2 6 1. 43 6 0. 21 0. 02 1 5. 57 0. 86 0 9. 79 0. 83 4 2. 30 0. 27 0 0. 11 0. 02 2 0. 02 0. 00 6 99 .3 7 lz c 84 7 81 6 0. 34 6 34 .1 3 8. 86 7 12 .3 5 5. 33 1 9. 67 2. 96 8 6. 76 2. 90 6 18 .7 4 5. 84 9 0. 26 0. 04 2 6. 68 0. 50 8 8. 68 2. 07 5 1. 68 0. 60 0 0. 09 0. 03 4 0. 02 0. 00 9 99 .0 6 lz c 80 8 79 8 0. 36 2 29 .4 9 11 .2 00 15 .3 1 6. 99 3 9. 42 3. 22 4 9. 07 4. 27 7 20 .3 8 6. 27 9 0. 26 0. 04 6 5. 87 0. 53 5 7. 20 2. 73 2 1. 58 0. 68 7 0. 09 0. 04 5 0. 02 0. 01 3 98 .7 0 lz c 78 4 74 2 0. 37 2 44 .6 6 6. 66 0 5. 07 4. 87 5 10 .4 5 4. 32 9 3. 92 1. 92 9 12 .5 2 2. 90 8 0. 22 0. 04 0 8. 37 3. 31 5 12 .0 9 3. 06 2 1. 86 0. 77 6 0. 13 0. 04 4 0. 05 0. 02 5 99 .3 4 lz c 72 3 70 3 0. 39 8 48 .7 0 2. 27 4 1. 61 0. 97 3 13 .9 8 1. 59 1 2. 01 1. 55 4 10 .0 1 2. 55 3 0. 19 0. 03 4 8. 28 2. 30 3 11 .7 3 1. 42 9 2. 46 0. 26 9 0. 18 0. 00 3 0. 06 0. 01 3 99 .1 9 lz b 70 3 68 1 0. 40 6 48 .7 6 1. 73 7 1. 73 0. 65 0 15 .6 0 2. 31 2 1. 77 1. 17 1 9. 89 1. 91 4 0. 17 0. 02 8 7. 50 2. 04 9 11 .1 9 0. 10 4 2. 73 0. 49 0 0. 18 0. 04 7 0. 05 0. 02 4 99 .5 7 lz b 63 4 58 0 0. 43 7 50 .4 6 0. 81 9 1. 32 0. 03 5 15 .4 5 3. 38 2 1. 27 0. 36 9 8. 44 1. 96 5 0. 17 0. 03 5 7. 13 2. 14 4 12 .3 2 1. 32 7 2. 64 0. 60 8 0. 20 0. 05 4 0. 08 0. 02 2 99 .4 9 lz b 55 8 48 8 0. 47 2 49 .0 3 1. 68 2 1. 31 0. 15 1 14 .8 3 1. 82 9 1. 71 0. 53 9 9. 77 2. 61 9 0. 18 0. 03 7 8. 64 2. 20 5 11 .2 1 1. 47 4 2. 45 0. 30 6 0. 18 0. 03 1 0. 08 0. 02 7 99 .3 9 lz b 44 7 36 7 0. 52 6 49 .2 8 2. 15 7 1. 27 0. 08 8 16 .0 2 2. 69 6 1. 65 0. 27 5 9. 50 3. 86 6 0. 17 0. 04 8 7. 81 3. 00 6 10 .9 9 1. 29 5 2. 65 0. 47 7 0. 22 0. 05 5 0. 10 0. 03 3 99 .6 5 lz b 34 6 22 1 0. 57 6 48 .2 2 2. 53 6 1. 10 0. 32 6 14 .7 2 3. 99 6 1. 21 0. 30 7 10 .7 2 5. 52 8 0. 18 0. 07 0 10 .0 5 3. 53 5 11 .0 3 1. 85 4 2. 26 0. 54 5 0. 18 0. 06 2 0. 07 0. 03 6 99 .7 3 lz b 17 7 17 3 0. 66 7 49 .4 6 1. 19 9 1. 33 0. 31 7 18 .1 7 4. 36 5 1. 53 0. 12 7 7. 79 2. 96 4 0. 14 0. 04 8 6. 93 3. 04 3 11 .3 3 1. 03 5 2. 79 0. 58 8 0. 25 0. 03 0 0. 11 0. 00 6 99 .8 4 lz a 16 1 13 7 0. 67 6 48 .6 0 0. 43 2 1. 30 0. 16 9 17 .6 2 0. 52 1 1. 63 0. 37 5 8. 88 0. 38 6 0. 16 0. 00 7 8. 10 0. 58 8 10 .3 1 0. 56 2 2. 70 0. 08 4 0. 26 0. 02 5 0. 12 0. 01 6 99 .6 7 lz a 12 5 10 7 0. 69 6 47 .3 8 1. 23 9 1. 53 0. 31 8 15 .9 0 2. 17 4 1. 43 0. 46 3 11 .0 8 1. 92 4 0. 18 0. 03 2 9. 98 1. 91 6 9. 16 0. 88 5 2. 46 0. 27 8 0. 26 0. 02 1 0. 14 0. 01 7 99 .4 9 lz a 96 27 0. 71 2 48 .7 5 0. 56 1 1. 42 0. 01 9 18 .6 9 1. 53 7 1. 35 0. 19 1 8. 40 0. 84 6 0. 15 0. 01 5 7. 02 1. 73 9 10 .8 4 0. 52 0 2. 73 0. 17 7 0. 26 0. 01 0 0. 13 0. 01 5 99 .7 5 lz a 7 0. 76 4 48 .1 6 1. 36 16 .9 1 0. 58 10 .0 9 0. 16 9. 01 10 .5 2 2. 41 0. 24 0. 13 99 .5 8 ca lc ul at ed a s av er ag es (a ve .) of th re e co ns ec ut iv e sa m pl es w ith 1 σ sd c al cu la te d fo r n > 2. f : m el t r em ai ni ng a ft er t eg ne r et a l. (2 00 9) , ( eq .) 1. s tr at ig ra ph ic h ei gh t i s gi ve n as h ig h an d lo w r ef er ri ng to th e in te rv al u se d fo r av er ag in g. c om pl et e da ta s et in s up pl em en ta ry f ile s 1. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 49 of 139 geusbulletin.org as suggested by the petrography (fig. 9). other large variations can be attributed to heterogeneity due to modal layering, particularly in the lzc. such variation highlights the problems associated with sampling highly heterogeneous rocks, such as layered gabbros. however, these consistent trends attest to some success with the sampling of the average modal compositions. table 14 average zone and subzone major and trace element compositions zone uzc uzb uza mz lzc lzb lza n ave. sd ave. sd ave. sd ave. sd ave. sd ave. sd ave. sd 21 24 21 25 12 19 12 major element concentrations (wt%) sio2 46.67 2.49 42.73 4.21 44.88 2.09 41.78 4.08 37.46 9.49 49.11 1.73 48.23 1.35 tio2 2.52 0.49 3.71 1.06 5.18 1.07 7.79 2.80 9.91 6.43 1.41 0.46 1.35 0.23 al2o3 7.98 1.44 10.57 0.97 12.28 1.03 11.81 2.87 10.56 3.01 15.30 2.27 17.54 2.58 fe2o3 4.76 1.97 3.71 0.84 3.92 1.23 4.43 1.84 6.13 3.30 1.58 0.74 1.34 0.41 feo 22.80 3.39 20.42 2.42 14.78 1.77 14.74 2.87 16.42 5.18 9.56 2.74 9.38 2.30 feo* 27.55 2.71 24.12 5.29 18.69 2.29 19.16 5.54 22.54 8.08 11.13 2.86 10.72 2.31 mno 0.51 0.27 0.37 0.08 0.26 0.02 0.24 0.04 0.24 0.04 0.17 0.04 0.16 0.03 mgo 0.21 0.26 3.08 1.54 5.41 0.75 6.17 1.32 7.02 2.24 8.14 2.24 8.50 2.81 cao 9.26 1.02 9.36 1.13 9.56 0.60 9.95 0.98 9.41 2.71 11.44 1.23 10.13 1.07 na2o 2.60 0.52 2.92 0.34 2.80 0.31 2.28 0.57 1.85 0.59 2.54 0.40 2.62 0.38 k2o 0.52 0.24 0.31 0.11 0.21 0.06 0.13 0.04 0.11 0.04 0.20 0.04 0.25 0.02 p2o5 0.68 0.19 1.87 0.94 0.07 0.04 0.04 0.01 0.03 0.02 0.08 0.03 0.13 0.02 loi 1.20 0.39 0.60 0.12 0.55 0.17 0.70 0.08 0.70 0.14 total 99.71 99.03 99.33 99.96 99.68 100.23 100.34 selected average trace element concentrations (ppm) sc 37 49 47 54 52 36 24 v 36 36 374 982 1410 346 227 cr 11 8 8 10 71 197 79 co 2.6 4 16 59 122 126 192 ni 13 78 91 82 89 54 59 cu 296 835 995 128 86 94 138 zn 224 144 109 96 95 78 83 rb 12 3.9 2.2 1.1 1.3 2.6 4.3 sr 400 282 285 226 198 277 298 y 136 40 13 8 8.54 13 17 zr 216 62 53 47 49 53 69 nb 50 9 8 9 7 4.6 8 ba 238 75 57 32 30 50 63 hf 5 1.7 1.5 1.3 1.4 1.5 1.9 ta 2.9 0.83 0.84 0.77 0.55 0.3 0.53 pb 2.6 1.1 0.63 0.91 0.90 1.1 1.3 th 1.5 0.51 0.22 0.10 0.13 0.33 0.56 u 0.40 0.15 0.07 0.03 0.04 0.09 0.16 la 28 13 3.3 1.8 2.1 4.3 7 yb 1.7 0.35 0.17 0.13 0.12 0.17 0.20 ls estimates by mcbirney (1989a) sio2 45.13 41.17 42.62 42.43 40.73 48.39 47.88 tio2 2.58 4.01 5.61 6.73 6.85 1.43 1.48 al2o3 7.71 9.37 11.05 11.43 10.92 12.43 16.65 fe2o3 5.69 5.63 4.98 6.07 6.63 2.56 2.35 feo 23.00 21.19 17.80 14.37 14.95 10.42 9.35 feo* 28.12 26.26 22.28 19.83 20.92 12.72 11.46 mno 0.64 0.41 0.30 0.26 0.26 0.21 0.17 mgo 0.38 3.36 5.56 6.19 6.19 10.04 8.78 cao 9.95 9.23 8.53 9.79 9.68 11.46 10.08 na2o 2.38 2.55 2.52 2.21 1.96 2.11 2.54 k2o 0.40 0.35 0.26 0.21 0.16 0.20 0.27 p2o5 0.83 1.85 0.21 0.08 0.04 0.09 0.11 loi 1.12 0.77 0.64 0.50 0.26 0.56 0.54 total 99.81 99.89 100.08 100.27 98.63 99.90 100.20 n: number of analyses used for each average (ave.) and 1σ sd for n > 2. loi: loss on ignitation at 950°c (not determined for uzb and uza) feo*: total iron as feo. ls: layered series. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 50 of 139 geusbulletin.org 4.5.2 trace elements trace element concentrations in the gabbros (figs 34, 35; summarised in tables 15, 16) vary, like the major elements, with mineral assembly, mineral modes and differentiation as a function of stratigraphic position. the trace elements can be divided into two groups, those that are dominantly partitioned into the primocryst minerals (included) or into the coexisting melt (excluded; mcbirney 1998, 2002). the former, with partition coefficients above unity, will occur in high concentrations in the early-formed gabbros. the latter, with partition coefficients below unity, will be enriched in the residual melt fraction with advanced differentiation and will thus be relatively depleted in the early gabbros. earlier studies of trace elements in the skaergaard intrusion include wager & mitchell (1951), haskin & haskin (1968), brooks (1969) and paster et al. (1974). 4.5.2.1 included trace elements the included trace elements are particularly well illustrated by the transition metals (fig. 34). these are highly to partially included elements with bulk-partitioning coefficients often well above unity that are known to be partitioned preferentially into olivine (ni and co), magnetite and ilmenite (sc, v, cr, co, ni and zn) and pyroxenes (sc, cr, co and ni; wager & mitchell 1951; paster et al. 1974; mcbirney 1998; jang & naslund 2003). the concentrations of both the included and excluded trace elements are shown in table 15. despite large scatter in some subzones, the variations for v, cr and ni show moderate to strong depletions with stratigraphic height reflecting their preferred partitioning into specific early primocrysts and subzones (fig. 34). vanadium is constant at 200–400 ppm in lza and lzb, followed by a marked enrichment in lzc to >2000 ppm due to the appearance of feti oxides as a primocryst. it is terminated by an asymptotic depletion from the base of lzc and into the uz approaching c. 5 ppm. the uz variation is caused by the concurrent depletion of v in the coexisting melt as well as the upward decreasing modal abundances, mainly of magnetite (fig. 15). chromium shows low concentrations in lza of 50–100 ppm, followed by an irregular increase in lzb to c. 300 ppm due to the appearance of clinopyroxene. subsequently, cr strongly depleted in the lzb and lzc, reaches low concentrations in the mz and approaches c. 9 ppm in the uz, again due to the concurrent strong depletion in the melt. nickel likewise decreases from c. 200 ppm at the base of lza to the top of the mz and further into the uzb, where it reaches low concentrations approaching 3 ppm with the consistent appearance of olivine and other mafic minerals throughout the stratigraphy and the concurrent depletion of ni in the coexisting melt. the high modal abundance of olivine at the base of uzb is not recorded by ni because of the low melt content of ni in the iron-rich melts of the uz. the prominent variations for these highly included elements are strongly depleted (figs 15, 34) due to early partitioning into olivine (ni), pyroxene (cr) and feti oxides (v), respectively. copper is consistently low at 100–150 ppm from the lz until the upper part of the mz where sulphides become saturated (andersen 2006; nielsen et al. 2015). because cu is strongly partitioned into sulphides (as minerals or immiscible melt), the concentrations in uza and uzb remain relatively constant (fig. 34). the variation of zn, with a systematic upward increase from c. 75 to 150 ppm and a spike in the upper part do not partition into the sulphides to any significant extent. the variations of co and sc differ from the other transition metals because they are partitioned into all mafic and feti oxide minerals. scandium concentration is constantly low in lza at c. 20 ppm, increases to a relatively constant concentration of about 30–40 ppm in lzb because of the appearance of clinopyroxene, but subsequently becomes highly variable, which is difficult to interpret in terms of the zones and subzones and modal 1000 2000 3000010 30 50 70 90 100 200 300 4000 50 100 150 2000 50 150 2500500 1000 1500 20000100 200 300 4000 0 200 400 600 800 1000 1200 1400 1600 1800 2000 2200 0 200 400 600 800 1000 1200 1400 1600 1800 2000 2200 vsc cr co zncuni 90-22 st ra ti gr ap hi c he ig ht (m ) transition trace elements (ppm) lza lzb lzc mz uza uzb uzc sh/ubs fig. 34 transition trace elements (sc, v, cr, co, ni, cu and zn; ppm) as a function of stratigraphic height (m) in the ls. blue horizontal dashed lines: divisions of the ls. abbreviations in fig. 2. red vertical dashed line: location of drill core 90-22. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 51 of 139 geusbulletin.org ta bl e 15 s um m ar y of g ab br o tr ac e el em en t c om po si tio ns zo ne   st ra ti g ra ph ic he ig ht (m ) sc (p pm ) v (p pm ) cr (p pm ) n i (p pm ) co (p pm ) cu (p pm ) zn (p pm ) rb (p pm ) sr (p pm ) y (p pm ) zr (p pm ) n b (p pm ) ba (p pm ) h f (p pm ) w (p pm ) ta (p pm ) pb (p pm ) th (p pm ) u (p pm ) hi gh lo w av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd u bs 21 69 21 65 31 3 25 1 22 3 3. 6 1. 3 15 11 14 3 15 5 24 8 54 21 2 47 4 11 5 15 7 32 30 2 12 93 44 40 7 48 8 0. 31 0. 42 0. 03 6 2. 17 0 4. 4 0. 3 2. 5 0. 3 0. 64 0. 06 u zc 21 65 21 65 44 24 22 3. 7 3. 4 72 1 25 9 16 57 0 11 7 39 8 52 34 9 9 0. 31 3. 0 3. 4 2. 1 0. 56 u zc 21 65 21 64 45 6 49 22 8 10 2. 5 0. 3 2. 5 0. 7 11 2 19 31 7 7 15 2 45 7 45 21 9 26 24 6 17 10 5 28 31 8 20 7 0. 36 0. 31 0. 02 6 1. 05 7 3. 5 0. 1 1. 9 0. 4 0. 51 0. 10 u zc 21 63 21 47 39 5 61 0 2. 2 1. 6 3. 1 0. 9 2. 6 1. 4 14 2 64 23 2 20 12 3 41 9 12 6 15 4 38 31 8 18 3 52 18 27 3 69 7 2. 82 0. 24 0. 04 2. 8 1. 06 2 2. 8 0. 5 1. 7 0. 4 0. 45 0. 07 u zc 21 44 21 41 37 61 3. 1 3. 5 4. 9 18 4 71 19 6 9 32 3 14 5 12 7 28 16 2 3. 4 0. 19 1. 6 2. 0 1. 1 0. 31 u zc 20 91 20 81 30 2 24 32 6 5 1. 5 1. 1 34 22 67 9 95 15 6 16 3. 8 1. 5 30 0 26 77 11 52 8 9 1 78 10 1. 4 0. 1 0. 83 0. 17 7 0. 92 0. 30 0. 50 0. 07 0. 14 0. 02 u zc 20 75 20 60 30 1 6 2 10 2 0. 64 0. 55 29 21 74 4 12 8 15 8 14 4. 2 1. 2 25 8 32 85 8 71 15 11 0 79 5 1. 9 0. 3 0. 93 0. 12 7 0. 78 0. 11 0. 59 0. 07 0. 18 0. 00 u zc 20 60 20 46 37 19 6 2 10 4 0. 90 1. 01 37 14 64 0 16 9 15 3 15 4. 6 0. 7 29 8 75 62 24 77 4 11 1 87 5 2. 0 0. 1 0. 85 0. 19 9 1. 0 0. 1 0. 66 0. 15 0. 19 0. 03 u zb 20 44 20 30 49 5 8 0. 23 51 36 0 56 15 9 6 28 4 50 91 13 99 2. 5 1. 1 1. 9 0. 82 0. 21 u zb 20 24 19 84 73 5 9 0 5 0 54 1 66 8 18 4 17 4 12 4. 6 0. 9 31 9 28 35 1 49 7 8 0 83 17 1. 5 0. 1 0. 75 0. 04 1. 1 0. 2 0. 50 0. 09 0. 12 0. 00 u zb 19 62 19 21 63 4 9 0 4. 4 0. 0 59 1 83 6 10 8 15 1 4 4. 6 1. 4 29 6 23 32 1 50 7 7 0 77 2 1. 5 0. 2 0. 65 0. 04 1. 1 0. 1 0. 47 0. 01 0. 13 0. 00 u zb 19 02 18 62 52 2 1. 0 0. 5 9 0 4. 4 0. 1 64 8 77 9 24 13 8 5 4. 9 1. 3 29 7 2 32 1 60 2 8 1 82 7 1. 6 0. 1 0. 67 0. 04 1. 0 0. 1 0. 52 0. 04 0. 15 0. 01 u zb 18 43 18 01 49 8 16 1 9 0 3. 5 0. 1 75 10 86 1 13 6 13 0 13 3. 6 1. 5 25 9 40 39 6 59 7 8 2 68 12 1. 6 0. 2 0. 76 0. 15 1. 1 0. 3 0. 50 0. 15 0. 17 0. 03 u zb 17 82 17 48 42 2 22 3 9 0 3. 6 0. 1 97 7 78 6 62 13 8 1 3. 9 0. 1 26 3 1 43 4 71 4 9 1 73 2 1. 9 0. 1 0. 96 0. 11 0. 90 0. 05 0. 52 0. 04 0. 17 0. 01 u zb 17 28 16 91 30 3 37 10 9 0 5 1 10 0 13 12 51 28 7 13 7 2 2. 5 0. 5 26 8 9 53 10 59 6 10 2 61 6 1. 6 0. 1 1. 0 0. 2 0. 71 0. 11 0. 43 0. 05 0. 16 0. 02 u zb 16 71 16 30 34 2 44 13 6 0 4. 4 0. 4 98 15 97 0 64 12 5 11 2. 7 0. 8 28 3 30 49 2 60 1 9 1 69 12 1. 6 0. 1 0. 87 0. 10 0. 93 0. 37 0. 44 0. 07 0. 16 0. 02 u zb 16 20 16 00 46 12 12 4 32 6 0 6 0 10 3 5 10 08 44 7 14 4 20 2. 5 0. 9 26 7 25 31 22 57 7 7 1 63 13 1. 6 0. 2 0. 67 0. 08 0. 79 0. 17 0. 35 0. 08 0. 12 0. 04 u za 15 61 15 27 44 0 10 3 34 6 0 6 0 87 3 79 7 35 1 11 6 7 2. 4 0. 7 32 2 12 12 2 51 11 10 4 71 10 1. 4 0. 3 0. 92 0. 33 0. 72 0. 16 0. 21 0. 07 0. 07 1 0. 01 6 u za 15 06 14 31 49 3 16 0 32 9 0 11 1 87 12 10 48 18 5 10 9 1 3. 5 3. 3 30 4 33 12 3 52 8 9 2 63 19 1. 4 0. 2 0. 92 0. 17 0. 56 0. 21 0. 20 0. 10 0. 05 2 0. 02 1 u za 14 03 13 43 47 3 27 0 72 9 0 12 0 89 8 92 8 61 10 7 11 1. 9 0. 1 30 1 24 12 0 49 9 7 3 55 5 1. 4 0. 2 0. 76 0. 28 0. 53 0. 02 0. 20 0. 02 0. 05 8 0. 00 4 u za 13 23 12 82 47 1 40 6 14 9 0 18 1 10 4 13 69 9 40 0 11 1 3 1. 6 0. 2 26 5 7 11 1 48 3 8 1 47 3 1. 3 0. 1 0. 87 0. 07 0. 47 0. 03 0. 19 0. 02 0. 05 7 0. 01 0 u za 12 68 12 37 47 2 57 1 30 8 8 0 21 5 95 14 14 19 49 0 11 4 23 1. 6 0. 8 26 1 24 13 5 49 11 7 2 51 12 1. 3 0. 3 0. 69 0. 15 0. 79 0. 60 0. 23 0. 17 0. 07 3 0. 02 9 u za 12 18 11 88 49 4 48 4 13 7 8 0 21 3 90 9 60 2 34 5 10 3 10 2. 5 0. 2 26 3 16 17 1 68 2 10 0 57 4 1. 8 0. 1 0. 91 0. 04 0. 61 0. 05 0. 31 0. 03 0. 09 1 0. 00 6 u za 11 78 11 31 44 8 62 3 17 7 7 2 24 4 81 29 14 74 68 9 10 5 19 2. 1 0. 4 27 9 46 12 5 57 8 9 2 54 4 1. 5 0. 2 0. 79 0. 07 0. 71 0. 26 0. 23 0. 06 0. 06 9 0. 01 8 m z 10 98 10 55 50 11 74 7 17 0 4. 2 0. 3 31 6 79 13 38 4 45 9 10 1 5 1. 6 0. 0 25 0 66 9 2 51 6 10 2 41 8 1. 4 0. 2 0. 15 0. 09 0. 77 0. 23 0. 99 0. 04 0. 15 0. 01 0. 04 3 0. 00 4 m z 10 53 10 32 64 59 9 5 33 78 74 36 89 0. 89 19 1 9 54 14 27 1. 5 0. 08 5 1. 2 0. 87 0. 07 5 0. 02 4 m z 10 24 10 09 56 3 89 2 21 6 4. 4 0. 4 53 7 85 9 10 3 3 99 5 1. 3 0. 4 22 0 7 10 3 44 1 7 2 34 5 1. 2 0. 0 0. 08 9 0. 00 7 0. 57 0. 16 0. 95 0. 07 0. 13 0. 05 0. 03 7 0. 01 4 m z 10 03 10 02 51 13 10 84 45 4 6 3 49 16 84 29 10 8 11 10 1 16 1. 1 0. 5 25 3 96 8 3 48 7 9 3 37 14 1. 3 0. 2 0. 09 4 0. 00 3 0. 70 0. 31 0. 94 0. 10 0. 12 0. 06 0. 03 2 0. 01 2 m z 98 8 96 1 65 12 10 91 25 4 4. 3 1. 0 70 10 98 11 10 9 23 10 3 8 1. 1 0. 2 15 1 49 12 6 60 6 12 3 25 4 1. 7 0. 2 0. 09 0 0. 00 2 0. 99 0. 28 0. 92 0. 06 0. 11 0. 04 0. 03 3 0. 01 0 m z 95 3 93 0 52 11 90 5 72 82 10 7 6 96 1. 2 22 8 8 48 9 33 1. 3 0. 09 1 0. 77 0. 89 0. 11 0. 03 3 m z 92 1 90 0 45 10 84 6 52 8 42 61 73 14 68 25 57 28 87 14 0. 82 0. 47 27 3 65 5 1 33 9 6 2 33 11 1. 0 0. 2 0. 08 9 0. 00 5 0. 48 0. 23 0. 86 0. 11 0. 06 0 0. 05 6 0. 01 9 0. 01 4 m z 89 0 87 5 48 5 14 09 24 1 7 2 95 18 84 11 78 17 95 7 0. 86 0. 40 24 1 30 5 1 40 5 8 1 29 5 1. 2 0. 1 0. 08 7 0. 00 5 0. 71 0. 11 0. 86 0. 05 0. 05 8 0. 02 4 0. 01 9 0. 01 0 lz c 84 7 81 6 53 13 73 6 99 93 95 20 96 0. 89 20 7 6 48 8 26 1. 4 1. 3 0. 65 0. 87 0. 08 6 0. 02 6 lz c 80 8 79 8 51 1 23 14 10 29 22 10 16 3 56 13 4 49 96 11 11 6 24 0. 73 0. 49 15 5 81 3. 1 3. 6 56 9 11 4 20 12 1. 6 0. 3 0. 08 7 0. 00 2 0. 93 0. 36 0. 84 0. 08 0. 06 7 0. 05 4 0. 02 4 0. 01 2 lz c 78 4 74 2 60 21 13 52 69 9 10 5 10 3 11 0 13 73 17 81 9 90 11 1. 4 0. 7 18 1 87 13 9 50 10 5 3 31 10 1. 5 0. 3 0. 48 6 0. 68 1 0. 39 0. 30 0. 88 0. 10 0. 16 0. 12 0. 04 7 0. 03 2 lz c 72 3 70 3 44 4 60 1 48 0 15 1 72 11 5 50 59 23 71 14 79 13 2. 2 0. 6 24 8 28 12 1 41 8 3. 3 0. 9 43 2 1. 2 0. 2 0. 87 0 0. 67 9 0. 22 0. 05 1. 0 0. 0 0. 22 0. 09 0. 05 7 0. 01 6 lz b 70 3 68 1 38 6 60 8 40 2 39 3 31 8 11 3 38 58 14 79 18 79 7 1. 7 0. 8 27 9 47 9 1 36 11 3. 2 0. 9 45 12 1. 1 0. 3 0. 09 7 0. 00 8 0. 21 0. 05 1. 0 0. 1 0. 19 0. 13 0. 05 1 0. 03 2 lz b 63 4 58 0 43 13 37 3 10 3 13 4 56 83 30 44 14 84 18 75 7 2. 5 0. 7 27 4 62 14 2 51 6 4. 3 0. 8 51 11 1. 4 0. 1 0. 24 1 0. 20 4 0. 28 0. 06 1. 1 0. 1 0. 29 0. 08 0. 08 2 0. 02 4 lz b 55 8 48 8 38 4 29 9 34 13 3 19 13 1 63 58 21 92 20 79 13 2. 5 0. 9 27 1 35 14 4 52 16 4. 2 1. 4 48 9 1. 4 0. 4 0. 70 7 0. 39 2 0. 30 0. 09 0. 98 0. 20 0. 36 0. 16 0. 09 1 0. 03 5 (c on tin ue d) https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 52 of 139 geusbulletin.org variations. cobalt similarly records variable concentrations until the base of uzb, where a modest increase may be related to the high modal olivine. interestingly, sc (30–70 ppm) and co (100–50 ppm) show contrasting variations from the base of uzb with a strong increase in sc and a corresponding decrease in co that are not easily explained by modal variations, but may suggest behaviour related to liquid immiscibility as in the uzb (figs 15, 34; see discussions in sections 5.7.7 and 5.12). 4.5.2.2 excluded trace elements the excluded group of elements does not always show systematic variations as a function of stratigraphic position directly relatable to modal content of major silicates and oxide minerals in the zones and subzones (table 15). the group of elements shown in fig. 35, plotted on a semi-log scale to easier allow pattern recognition, is preferentially excluded from the solid fraction during crystallisation and thus resides principally in the coexisting melts (including trapped melt). an element is defined as excluded if its bulk partition coefficient is below unity (mcbirney 2002). examples are rb, sr, y, ba, the rare-earth elements (ree; table 16) and u that are mostly excluded throughout the ls, although some of these may have partitioning coefficients above unity. an example of the latter is the ree that are excluded in the lza to uza, but intermittently or nearly included in uzb, when apatite first appears as a primocryst (figs 35, 36). concentrations of excluded trace elements are shown in tables 15 and 16. see also discussion of bulk-partitioning coefficients in section 5.11.3. observable patterns in trace element variations correlate with stratigraphic height and subzone mineralogy (fig. 35). most of the elements show a systematic upward decrease in the lza to lzc, exemplified by rb that varies from 4 ppm at the base of lza to 2 ppm in the upper part of lzc. the systematic decrease in these elements is interpreted to record the concurrent decrease in the trapped liquid content (e.g. henderson 1970, 1975; tegner et al. 2009) that appears to mask any expected upward increases in the coexisting residual magma due to the low bulk partition of these elements. the correlations between the trapped melt content versus rb and la in the lz are illustrated in fig. 37 according to estimates of trapped melt content made by tegner et  al. (2009). the effect of feti oxide crystallisation in the lzc can be seen in most of the excluded elements as decreasing concentrations (rb, sr, y, ba, la and other ree, th and u). this observation can be attributed to the low concentrations of these elements and the high modal amount (<20–30 wt%) of feti oxides in lzc (fig. 15) and the low trapped-melt content in lzc documented by tegner et al. (2009), but may also be related to the observed ta bl e 15 (c on tin ue d) s um m ar y of g ab br o tr ac e el em en t c om po si tio ns zo ne   st ra ti g ra ph ic he ig ht (m ) sc (p pm ) v (p pm ) cr (p pm ) n i (p pm ) co (p pm ) cu (p pm ) zn (p pm ) rb (p pm ) sr (p pm ) y (p pm ) zr (p pm ) n b (p pm ) ba (p pm ) h f (p pm ) w (p pm ) ta (p pm ) pb (p pm ) th (p pm ) u (p pm ) hi gh lo w av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd lz b 44 7 36 7 34 1 29 0 31 11 7 8 13 0 83 55 29 10 9 10 83 16 2. 8 1. 0 28 9 49 15 5 61 17 5 1 56 15 1. 6 0. 4 0. 11 7 0. 01 8 0. 32 0. 06 1. 2 0. 1 0. 37 0. 18 0. 10 0. 05 lz b 34 6 22 1 35 1 26 0 23 23 6 61 18 7 92 69 36 88 28 82 19 2. 5 1. 8 25 0 58 12 6 57 40 4. 8 3. 3 43 15 1. 5 0. 9 0. 11 2 0. 03 1 0. 30 0. 20 1. 1 0. 2 0. 33 0. 30 0. 10 0. 10 lz b 17 7 17 3 28 11 24 5 80 16 9 13 3 11 4 50 43 17 11 2 9 71 8 3. 6 0. 1 29 9 82 13 3 63 13 6 1 59 10 1. 7 0. 3 0. 42 0 0. 51 5 0. 40 0. 05 1. 2 0. 0 0. 42 0. 01 0. 12 0. 01 lz a 16 1 13 7 24 1 21 1 2 98 31 16 3 23 56 6 12 2 4 79 3 4. 1 0. 9 30 0 14 15 2 68 4 7 1 63 3 1. 8 0. 1 0. 12 9 0. 00 9 0. 45 0. 08 1. 3 0. 1 0. 46 0. 10 0. 13 0. 03 lz a 12 5 10 7 23 1 22 7 26 55 12 22 7 61 70 16 15 7 44 89 13 4. 5 0. 5 27 8 36 18 2 73 4 8 1 63 3 2. 0 0. 1 0. 13 8 0. 00 5 0. 54 0. 11 1. 3 0. 0 0. 56 0. 06 0. 16 0. 01 lz a 96 27 24 2 23 0 9 77 13 15 2 66 49 12 14 2 5 80 6 4. 1 0. 2 31 6 21 18 2 72 15 8 0 65 3 1. 9 0. 3 0. 13 7 0. 01 4 0. 51 0. 05 1. 5 0. 3 0. 57 0. 14 0. 17 0. 04 lz a 7 26 24 1 85 22 7 63 13 1 83 4. 6 29 6 16 62 9 61 1. 8 0. 17 8 0. 63 1. 3 0. 67 0. 15 ca lc ul at ed a s av er ag es (a ve .) of 3 c on se cu tiv e sa m pl es w ith 1 σ sd fo r n > 2 . c u w as a na ly se d us in g xr f, w hi le th e re m ai ni ng w er e by ic pm s. m is si ng d at a re fle ct s ei th er n ot a na ly se d or b el ow d et ec tio n lim its . r es ul ts fo r w a re no t g iv en fo r t he 9 022 d ril l c or e si nc e th es e sa m pl es w as p re pa re d in a w ol fr am c ar bi de m or te r. st ra tig ra ph ic h ei gh t i s gi ve n as h ig h an d lo w re fe rr in g to th e in te rv al u se d fo r a ve ra gi ng . c om pl et e da ta s et in s up pl em en ta ry f ile s1 . https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 53 of 139 geusbulletin.org ta bl e 16 s um m ar y of g ab br o re e co m po si tio ns zo ne st ra tig ra ph ic he ig ht (m ) la (p pm ) ce (p pm ) pr (p pm ) n d (p pm ) sm (p pm ) eu (p pm ) g d (p pm ) tb (p pm ) d y (p pm ) h o (p pm ) er (p pm ) tm (p pm ) yb (p pm ) lu (p pm ) hi gh lo w av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd u bs 21 69 21 65 56 8 15 2 28 24 6 11 7 32 31 9 11 4 29 8 4. 8 1. 3 30 7 5 1 14 3 1. 9 0. 4 12 3 1. 8 0. 4 u zc 21 65 21 65 35 98 16 82 23 12 22 3. 6 22 4. 0 11 1. 6 11 1. 7 u zc 21 65 21 64 57 5 15 8 13 26 2 13 2 11 37 3 13 1 35 2 6 1 39 3 7 1 19 2 2. 9 0. 4 18 2 2. 9 0. 3 u zc 21 63 21 47 37 11 10 3 28 17 4 85 18 24 5 9 2 23 4 4. 0 0. 8 25 6 4. 7 1. 1 13 3 2. 0 0. 5 13 3 2. 1 0. 4 u zc 21 44 21 41 29 80 13 66 19 6 19 3. 3 21 4. 1 12 1. 8 12 1. 9 u zc 20 91 20 81 12 1 33 3 6 0 31 2 9 1 4. 0 0. 3 11 1 1. 8 0. 1 12 1 2. 4 0. 2 7 1 1. 1 0. 1 7 1 1. 2 0. 1 u zc 20 75 20 60 14 0 39 0 6 0 34 1 10 0 3. 8 0. 3 12 0 2. 0 0. 1 13 1 2. 6 0. 0 7 0 1. 1 0. 0 8 0 1. 3 0. 1 u zc 20 60 20 46 13 2 35 7 6 1 30 7 9 2 3. 5 0. 1 10 3 1. 6 0. 5 10 4 2. 0 0. 8 6 2 0. 83 0. 40 5 3 0. 88 0. 45 u zb 20 44 20 30 15 41 7 34 9 3. 8 10 1. 5 9 1. 7 4. 5 0. 63 3. 8 0. 62 u zb 20 24 19 84 11 0 31 1 5 0 27 2 7 0 3. 5 0. 2 8 1 1. 2 0. 0 7 0 1. 2 0. 0 2. 9 0. 1 0. 39 0. 02 2. 3 0. 2 0. 36 0. 04 u zb 19 62 19 21 10 0 29 1 4. 8 0 26 1 7 0 3. 0 0. 1 8 0 1. 1 0. 0 6 0 1. 1 0. 0 2. 7 0. 1 0. 34 0. 02 2. 0 0. 1 0. 30 0. 02 u zb 19 02 18 62 11 0 31 1 5. 0 0 27 0 7 0 3. 0 0. 1 8 0 1. 1 0. 0 6 0 1. 1 0. 0 2. 6 0. 1 0. 33 0. 01 1. 9 0. 1 0. 27 0. 01 u zb 18 43 18 01 13 2 37 5 6 1 32 4 9 1 3. 1 0. 4 9 1 1. 3 0. 2 7 1 1. 3 0. 2 3. 1 0. 4 0. 38 0. 05 2. 2 0. 2 0. 31 0. 03 u zb 17 82 17 48 14 1 39 4 6 1 34 4 9 1 3. 3 0. 2 10 1 1. 5 0. 1 8 1 1. 4 0. 1 3. 4 0. 3 0. 42 0. 03 2. 4 0. 2 0. 33 0. 02 u zb 17 28 16 91 17 3 50 10 8 2 43 9 12 2 3. 7 0. 6 13 3 1. 8 0. 4 10 2 1. 7 0. 3 4. 0 0. 8 0. 48 0. 08 2. 6 0. 3 0. 35 0. 04 u zb 16 71 16 30 16 1 45 2 7 0 39 2 11 0 3. 4 0. 2 12 0 1. 7 0. 1 9 0 1. 6 0. 1 3. 8 0. 2 0. 45 0. 03 2. 5 0. 1 0. 34 0. 02 u zb 16 20 16 00 9 8 26 23 4. 1 3. 8 22 21 6 6 2. 4 1. 2 7 6 1. 0 0. 8 6 4 1. 0 0. 7 2. 5 1. 5 0. 32 0. 16 1. 9 0. 7 0. 26 0. 09 u za 15 61 15 27 3. 2 0. 9 8 2 1. 3 0. 3 6 1 1. 9 0. 3 1. 7 0. 1 2. 4 0. 3 0. 39 0. 06 2. 2 0. 3 0. 48 0. 06 1. 2 0. 2 0. 16 0. 03 1. 1 0. 2 0. 17 0. 02 u za 15 06 14 31 3. 0 1. 1 8 3 1. 2 0. 4 6 2 2. 0 0. 5 1. 5 0. 2 2. 4 0. 5 0. 39 0. 08 2. 3 0. 5 0. 49 0. 09 1. 3 0. 2 0. 18 0. 03 1. 2 0. 2 0. 17 0. 02 u za 14 03 13 43 3. 0 0. 1 8 0 1. 2 0. 0 6 0 1. 9 0. 0 1. 4 0. 1 2. 3 0. 0 0. 38 0. 00 2. 2 0. 0 0. 47 0. 01 1. 2 0. 0 0. 17 0. 00 1. 1 0. 0 0. 16 0. 00 u za 13 23 12 82 2. 8 0. 3 7 1 1. 1 0. 1 6 0 1. 8 0. 1 1. 2 0. 0 2. 3 0. 2 0. 40 0. 10 2. 1 0. 2 0. 46 0. 03 1. 2 0. 1 0. 17 0. 01 1. 1 0. 1 0. 16 0. 01 u za 12 68 12 37 3. 2 1. 6 8 4 1. 3 0. 6 7 3 2. 0 0. 7 1. 3 0. 2 2. 6 0. 7 0. 52 0. 12 2. 5 0. 8 0. 52 0. 14 1. 3 0. 4 0. 19 0. 06 1. 2 0. 3 0. 18 0. 05 u za 12 18 11 88 4. 3 0. 2 11 0 1. 7 0. 1 8 0 2. 5 0. 0 1. 4 0. 0 3. 1 0. 1 0. 56 0. 02 3. 0 0. 1 0. 62 0. 02 1. 6 0. 1 0. 23 0. 01 1. 4 0. 0 0. 21 0. 01 u za 11 78 11 31 3. 3 0. 6 8 2 1. 3 0. 3 7 1 1. 9 0. 5 1. 2 0. 1 2. 5 0. 6 0. 43 0. 14 2. 2 0. 8 0. 47 0. 15 1. 2 0. 4 0. 17 0. 06 1. 1 0. 4 0. 16 0. 05 m z 10 98 10 55 2. 3 0. 2 6 0 0. 96 0. 05 5 0 1. 6 0. 2 1. 0 0. 1 2. 1 0. 2 0. 33 0. 05 1. 8 0. 3 0. 38 0. 07 1. 0 0. 2 0. 13 0. 03 0. 87 0. 18 0. 13 0. 03 m z 10 53 10 32 1. 5 4. 1 0. 71 4. 1 1. 4 0. 85 2. 0 0. 32 1. 8 0. 39 1. 0 0. 13 0. 91 0. 14 m z 10 24 10 09 2. 1 0. 6 6 2 0. 94 0. 25 5 1 1. 6 0. 3 0. 98 0. 10 2. 2 0. 3 0. 35 0. 06 2. 0 0. 4 0. 42 0. 08 1. 1 0. 2 0. 14 0. 03 0. 96 0. 18 0. 15 0. 03 m z 10 03 10 02 1. 9 0. 7 5 2 0. 82 0. 23 4. 5 1. 1 1. 4 0. 4 0. 96 0. 20 2. 0 0. 4 0. 30 0. 08 1. 6 0. 5 0. 35 0. 10 0. 90 0. 29 0. 11 0. 04 0. 78 0. 25 0. 12 0. 04 m z 98 8 96 1 1. 8 0. 4 5 1 0. 92 0. 27 5 2 1. 8 0. 6 0. 84 0. 06 2. 4 0. 7 0. 38 0. 13 2. 2 0. 9 0. 46 0. 17 1. 2 0. 5 0. 16 0. 07 1. 1 0. 4 0. 16 0. 05 m z 95 3 93 0 1. 9 5 0. 83 4. 6 1. 4 0. 88 2. 0 0. 30 1. 7 0. 35 0. 91 0. 12 0. 78 0. 12 m z 92 1 90 0 1. 4 0. 8 3. 6 1. 9 0. 60 0. 26 3. 5 1. 1 1. 1 0. 2 0. 88 0. 19 1. 6 0. 2 0. 24 0. 02 1. 2 0. 1 0. 28 0. 02 0. 69 0. 05 0. 08 2 0. 00 6 0. 59 0. 02 0. 09 2 0. 00 5 m z 89 0 87 5 1. 2 0. 3 3. 3 0. 6 0. 56 0. 07 3. 3 0. 2 1. 0 0. 0 0. 77 0. 06 1. 6 0. 1 0. 23 0. 01 1. 2 0. 1 0. 27 0. 02 0. 67 0. 05 0. 07 9 0. 00 8 0. 58 0. 04 0. 08 9 0. 00 6 lz c 84 7 81 6 1. 5 0. 5 4. 1 1. 3 0. 68 0. 19 3. 8 0. 9 1. 2 0. 3 0. 73 0. 07 1. 7 0. 2 0. 26 0. 05 1. 4 0. 3 0. 31 0. 05 0. 78 0. 16 0. 09 5 0. 02 1 0. 67 0. 11 0. 10 0. 01 lz c 80 8 79 8 1. 1 0. 8 3. 1 2. 0 0. 50 0. 31 2. 8 1. 5 0. 87 0. 45 0. 55 0. 23 1. 4 0. 5 0. 19 0. 09 0. 87 0. 57 0. 21 0. 11 0. 53 0. 27 0. 06 1 0. 03 9 0. 47 0. 21 0. 07 6 0. 02 8 lz c 78 4 74 2 2. 5 1. 3 7 4 1. 1 0. 6 6 3 2. 0 1. 0 0. 85 0. 19 2. 4 1. 2 0. 41 0. 22 2. 4 1. 5 0. 49 0. 28 1. 3 0. 7 0. 17 0. 11 1. 1 0. 6 0. 17 0. 09 lz c 72 3 70 3 3. 0 0. 5 8 1 1. 2 0. 2 6 1 1. 9 0. 2 0. 96 0. 08 2. 3 0. 2 0. 38 0. 02 2. 1 0. 2 0. 44 0. 03 1. 2 0. 1 0. 15 0. 01 1. 0 0. 1 0. 15 0. 01 lz b 70 3 68 1 2. 8 1. 2 7 3 1. 1 0. 4 5 2 1. 7 0. 3 0. 94 0. 13 1. 9 0. 3 0. 33 0. 04 1. 7 0. 3 0. 37 0. 04 0. 97 0. 11 0. 12 0. 02 0. 84 0. 11 0. 12 0. 01 lz b 63 4 58 0 4. 2 0. 9 11 2 1. 7 0. 2 8 1 2. 4 0. 1 1. 1 0. 1 2. 8 0. 1 0. 46 0. 03 2. 7 0. 3 0. 54 0. 05 1. 4 0. 1 0. 19 0. 02 1. 2 0. 1 0. 18 0. 02 lz b 55 8 48 8 4. 2 1. 5 11 4 1. 6 0. 6 8 3 2. 3 0. 6 1. 1 0. 2 2. 8 0. 6 0. 45 0. 10 2. 6 0. 7 0. 53 0. 13 1. 4 0. 3 0. 21 0. 05 1. 1 0. 3 0. 18 0. 04 (c on tin ue d) https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 54 of 139 geusbulletin.org reduction in grain size (fig. 5). in contrast, the nb and ta concentrations are positively affected by feti oxide crystallisation in lzc, where nb increases from about 3 ppm to 8 ppm and ta increases from 0.3 ppm to 0.9 ppm, reflecting partitioning of these elements into ilmenite and to a lesser extent magnetite. one group of elements is approximately constant throughout the main part of the upper stratigraphy (uza to uzb): c. 50–60 ppm zr, c. 8–12 ppm nb, c. 1–2 ppm hf and c. 0.61–1 ppm ta. yttrium and lanthanum (and the other ree) abruptly increase in uzb (fig. 36) related to the appearance of apatite (fig. 14) and followed by a modest upward decrease. the sm/yb ratio shows marked increase upward in the uzb (fig. 36, normalised to typical mid-ocean ridge basalt (morb) concentration). ree patterns are further illustrated in detail in fig. 38 as normalised variations showing increasing steps in concentrations narrowly located at the uza–uzb boundary both for the light (lan) and heavy (ybn) ree. the uzb gabbros transgress the uzc gabbros (fig. 38) due to the change in slopes at the uza–uzb boundary (fig. 36). the systematically gently sloping patterns for the normalised ree are similarly consistent in the three individual uz subzones, but with positive eu-anomalies only for uza gabbros (fig. 38a), probably related to the lower plagioclase content in uzb and uzc compared to uza. a small group of elements (rb, sr and ba) suggests systematic, but small, increases upward with some noticeable exceptions, such as a small reduction for rb at the base of uza and in sr at the base of uzb, the latter can be related to high olivine and low plagioclase contents commencing at the base of this zone. finally, and unexplainably, there is a drop in pb, th and u in the central part of uza (fig. 35). most excluded element variations are capped by increasing and high concentrations in the uzc approaching the sh (fig. 35). for example, 10–20 ppm rb, 320–570 ppm sr, 150–220 ppm y, 120–400 ppm zr and 160–350 ppm ba. this is likely related to extreme differentiation of the residual magma and to the high amount of granophyric melt in uzc (figs 13, 14). for zr, the concentration has been sufficiently high in the melt to precipitate “rather abundant” zircon near the top of the uzc, as noted by wager & brown (1967). in summary, three important variables control excluded element behaviour in the gabbros as a function of stratigraphy, as follows: (1) trapped melt content partly controls the behaviour of the excluded elements with low bulk partition coefficients. this is particularly obvious in the upward decrease in several elements in the lza and lzb that is positively correlated with the trapped melt content. another possible effect of high melt content is in the upper part of the uzc where the trapped melt contributes to the bulk ta bl e 16 (c on tin ue d) s um m ar y of g ab br o re e co m po si tio ns zo ne st ra tig ra ph ic he ig ht (m ) la (p pm ) ce (p pm ) pr (p pm ) n d (p pm ) sm (p pm ) eu (p pm ) g d (p pm ) tb (p pm ) d y (p pm ) h o (p pm ) er (p pm ) tm (p pm ) yb (p pm ) lu (p pm ) hi gh lo w av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd av e. sd lz b 44 7 36 7 5 2 13 5 1. 9 0. 7 9 3 2. 6 0. 9 1. 2 0. 2 2. 9 0. 9 0. 48 0. 14 2. 8 0. 9 0. 56 0. 16 1. 5 0. 4 0. 20 0. 06 1. 3 0. 3 0. 19 0. 04 lz b 34 6 22 1 4. 1 2. 7 10 7 1. 6 0. 9 7 4 2. 1 1. 0 0. 94 0. 21 2. 4 1. 0 0. 39 0. 16 2. 2 1. 0 0. 45 0. 19 1. 2 0. 5 0. 16 0. 07 1. 1 0. 5 0. 15 0. 06 lz b 17 7 17 3 5 0 13 1 2. 0 0. 2 9 1 2. 5 0. 4 1. 1 0. 1 2. 7 0. 5 0. 45 0. 08 2. 5 0. 6 0. 50 0. 11 1. 3 0. 3 0. 18 0. 04 1. 1 0. 2 0. 17 0. 03 lz a 16 1 13 7 6 1 15 2 2. 3 0. 2 11 1 2. 8 0. 3 1. 2 0. 0 3. 1 0. 3 0. 49 0. 04 2. 8 0. 3 0. 56 0. 05 1. 5 0. 1 0. 20 0. 02 1. 3 0. 1 0. 19 0. 02 lz a 12 5 10 7 7 1 18 1 2. 6 0. 2 12 1 3. 2 0. 3 1. 2 0. 0 3. 5 0. 2 0. 56 0. 04 3. 2 0. 3 0. 63 0. 05 1. 7 0. 1 0. 23 0. 02 1. 5 0. 2 0. 21 0. 02 lz a 96 27 7 1 17 1 2. 5 0. 2 12 1 3. 1 0. 3 1. 3 0. 1 3. 4 0. 3 0. 54 0. 05 3. 1 0. 3 0. 63 0. 05 1. 7 0. 1 0. 23 0. 02 1. 4 0. 1 0. 21 0. 02 lz a 7 7 16 2. 4 11 3. 0 1. 2 3. 2 0. 51 2. 9 0. 58 1. 5 0. 21 1. 3 0. 19 ca lc ul at ed a s av er ag es (a ve .) of th re e co ns ec ut iv e sa m pl es w ith 1 σ sd fo r n > 2. m is si ng d at a re fle ct s ei th er n ot a na ly se d or r es ul ts b el ow d et ec tio n lim its . s tr at ig ra ph ic h ei gh t i s gi ve n as h ig h an d lo w r ef er ri ng to th e in te rv al u se d fo r av er ag in g. c om pl et e da ta s et in s up pl em en ta ry f ile s 1. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 55 of 139 geusbulletin.org partition coefficient by a melt-partition coefficient of unity, thus higher trapped melt content leads to higher contents of excluded elements. (2) the modal content of specific zones and subzones affects the trace element content by either increasing or decreasing the concentrations dependent on the specific partitioning coefficients. examples of such strong effects on the trace element concentrations are the appearances of apatite in uzb and feti oxides in the lzc. in the latter, high partition results in an increase (ta and nb) and low partition coefficients result in a decrease (rb, ba and ree). (3) fractional crystallisation results in a systematic upward increase in the main magma chamber for elements with low partition and reversely a decrease of elements with high partitioning coefficients. crystallisation and trapping of melt from such an evolving main chamber result in systematic variation in the solid fraction. depending on the bulk partition coefficients, the variation in the gabbros as a function of stratigraphy will vary so that low partition will result in increasing concentrations, unity partition (c. 1) will result in constant variations, and finally, high partition will result in decreasing variations upward in the stratigraphy. examples of this are several elements (zr, nb, hf and ta) that suggest near constant variations upward from lzc through uzc, while other elements (rb and ba) show a steady increase upward in the same stratigraphic intervals. the included elements v, cr and ni are examples of elements with high partition coefficients that decrease upward with stratigraphic position dependent of the magnitude of the partition coefficients. 4.5.3 co-variation of mineral modes and trace element concentrations the pearson correlation coefficients between the calculated modal mineralogy and trace element concentrations are shown in table 17 for both the lower (lz–uza) and the upper (uzb–uzc) ls. in the upper part of the ls, there is a remarkably high positive interdependency between all the excluded elements (including zn). the same group of elements is negatively correlated with f. this shows that the excluded elements behave in a similar fashion being partially excluded from the solid fraction and increasing upward in the melt stratigraphy (towards lower values of f and t). the same group of excluded trace elements in the same stratigraphic section is negatively correlated with the modal amount of st ra ti gr ap hi c he ig ht (m ) lza lzb lzc mz uza uzb uzc sh/ubs 90-22 st ra ti gr ap hi c he ig ht (m ) trace elements (ppm) lza lzb lzc mz uza uzb uzc sh/ubs 90-22 0 200 400 600 800 1000 1200 1400 1600 1800 2000 2200 0.1 1 10 100 100 1000 1 10 100 1000 10 100 1000 1 10 100 1000 0.1 1 10 0.1 1 10 100 0.1 1 10 10 100 1000 0 200 400 600 800 1000 1200 1400 1600 1800 2000 2200 0.01 0.1 1 10 0.01 0.1 10.1 1 10 100 0 200 400 600 800 1000 1200 1400 1600 1800 2000 2200 0 200 400 600 800 1000 1200 1400 1600 1800 2000 2200 rb sr y zr nb hf ba ta pb th ula fig. 35 selected alkali metal (rb), alkaline earth (sr, ba), high field strength (y, zr, nb, hf), rare-earth (la) and heavy (pb, th, u) trace elements (ppm) as a function of stratigraphic height (m) in the ls. plotted on a log scale. blue horizontal dashed lines: divisions of the ls. abbreviations divisions in fig. 2. red vertical dashed line: location of drill core 90-22. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 56 of 139 geusbulletin.org 0.1 1 10 1 10 0 200 400 600 800 1000 1200 1400 1600 1800 2000 2200 0.1 1 10 100 lan 0 200 400 600 800 1000 1200 1400 1600 1800 2000 2200 st ra ti gr ap hi c he ig ht (m ) ree elements normalised to morb (ppm) ybn sm/yb 90-22 lza lzb lzc mz uza uzb uzc sh/ubs fig. 36 selected rare earth element (lan and ybn) concentrations (ppm) and sm/yb ratio normalised to normal morb composition (sun & mcdonough 1989) as a function of stratigraphic height (m) in the ls. plotted on a log scale. blue horizontal dashed lines: divisions of the ls. abbreviations in fig. 2. red vertical dashed line: location of drill core 90-22. ree: rare earth elements. morb: mid-ocean ridge basalt. n: normalised to normal morb. trapped melt fraction tr ac e el em en t c on ce nt ra ti on (p pm ) lz rb la 0 1 2 3 4 5 6 7 8 0 0.1 0.2 0.3 0.4 1090 1100 1110 1120 1130 1140 t (°c) fig. 37 selected trace elements (rb and la) concentrations (ppm) as a function of the trapped melt content in the ls (tegner et al. 2009). corresponding melt temperature (t) is shown at the top of graph (thy et al. 2009b, 2013). ce nd sm eu gd dy ho er tm yb lula 0.01 0.1 1 10 100 re e co nc en tr at io ns n or m al is ed to m o rb uzc uza uzb 0.01 0.1 1 10 100 re e co nc en tr at io ns n or m al is ed to m o rb uz inverted melt compositions ce nd sm eu gd dy ho er tm yb lula uz gabbro compositions uzb uza uzc uzc uza uzb b a fig. 38 rare earth elements (ree) of uz normalised to normal morb (log scale; sun & mcdonough, 1989). a: gabbro compositions. b: inverted melt compositions. data are grouped as: uzc, uzb and uza. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 57 of 139 geusbulletin.org ta bl e 17 p ea rs on p ro du ct m om en t c or re la tio n co effi ci en t ( r) m at ri ce s fo r m in er al m od es a nd s el ec te d tr ac e el em en ts   su bz on es u zc a nd u zb f o l o px cp x pl m t il ap q o r co sc v cr n i zn rb sr y zr n b ba la ce n d sm eu g d d y h o er tm yb lu h f ta pb th u f 1. 00 o l –0 .0 3 1. 00 o px cp x –0 .5 5 –0 .6 2 1. 00 pl 0. 56 0. 17 –0 .7 0 1. 00 m t 0. 87 –0 .1 8 –0 .4 1 0. 38 1. 00 il 0. 83 0. 34 –0 .5 8 0. 33 0. 71 1. 00 ap 0. 81 0. 44 –0 .7 2 0. 43 0. 66 0. 93 1. 00 q –0 .5 8 –0 .1 5 0. 18 –0 .4 6 –0 .4 6 –0 .6 4 –0 .5 6 1. 00 o r –0 .6 4 –0 .1 2 0. 34 –0 .6 6 –0 .4 4 –0 .5 2 –0 .5 2 0. 74 1. 00 co 0. 94 0. 13 –0 .6 3 0. 70 0. 75 0. 80 0. 80 –0 .6 3– 0. 82 1. 00 sc –0 .0 6 0. 11 –0 .2 3 0. 48 –0 .0 7 –0 .2 4 –0 .2 0– 0. 04 –0 .1 5 0. 07 1. 00 v 0. 39 –0 .5 7 0. 21 –0 .3 1 0. 54 0. 25 0. 10 0. 08 0. 26 0. 10 –0 .2 8 1. 00 cr –0 .1 4 0. 11 –0 .2 0 0. 29 –0 .1 3 –0 .2 2 –0 .1 4 0. 06 0. 16 –0 .1 0 0. 13 –0 .3 9 1. 00 n i 0. 62 0. 02 –0 .4 4 0. 43 0. 62 0. 54 0. 52 –0 .4 3– 0. 16 0. 50 0. 35 0. 46 –0 .0 3 1. 00 zn –0 .6 4 –0 .1 3 0. 35 –0 .5 9 –0 .4 6 –0 .5 2 –0 .5 5 0. 66 0. 94 –0 .7 8 0. 03 0. 18 0. 19 –0 .1 4 1. 00 rb –0 .6 7 –0 .0 7 0. 25 –0 .5 0 –0 .5 0 –0 .5 8 –0 .5 6 0. 73 0. 97 –0 .8 1 0. 01 0. 12 0. 28 –0 .1 5 0. 95 1. 00 sr –0 .5 6 –0 .1 4 0. 18 –0 .3 2 –0 .4 0 –0 .5 3 –0 .4 8 0. 59 0. 88 –0 .6 9 0. 06 0. 10 0. 48 0. 00 0. 88 0. 94 1. 00 y –0 .6 2 –0 .0 8 0. 44 –0 .8 2 –0 .4 4 –0 .3 9 –0 .4 4 0. 62 0. 92 –0 .7 9– 0. 29 0. 28 –0 .1 0– 0. 27 0. 90 0. 85 0. 69 1. 00 zr –0 .5 2 –0 .1 3 0. 17 –0 .4 0 –0 .3 6 –0 .4 8 –0 .4 5 0. 67 0. 91 –0 .6 9– 0. 10 0. 19 0. 37 –0 .0 7 0. 84 0. 94 0. 94 0. 73 1. 00 n b –0 .5 1 –0 .0 7 0. 23 –0 .5 6 –0 .3 4 –0 .3 5 –0 .3 8 0. 62 0. 90 –0 .6 7– 0. 06 0. 24 0. 08 –0 .1 2 0. 96 0. 89 0. 79 0. 92 0. 78 1. 00 ba –0 .5 9 –0 .1 1 0. 23 –0 .4 9 –0 .4 2 –0 .5 1 –0 .4 9 0. 69 0. 97 –0 .7 6– 0. 04 0. 19 0. 30 –0 .0 9 0. 94 0. 99 0. 96 0. 84 0. 97 0. 90 1. 00 la –0 .5 0 0. 01 0. 19 –0 .6 2 –0 .3 3 –0 .3 1 –0 .3 2 0. 65 0. 94 –0 .6 8– 0. 15 0. 27 0. 04 –0 .1 0 0. 94 0. 91 0. 79 0. 95 0. 81 0. 98 0. 91 1. 00 ce –0 .4 9 0. 02 0. 18 –0 .6 2 –0 .3 3 –0 .3 0 –0 .3 0 0. 64 0. 93 –0 .6 7– 0. 16 0. 27 0. 04 –0 .1 0 0. 94 0. 90 0. 79 0. 95 0. 81 0. 98 0. 91 1. 00 1. 00 n d –0 .4 9 0. 02 0. 18 –0 .6 2 –0 .3 3 –0 .2 8 –0 .2 9 0. 63 0. 93 –0 .6 6– 0. 17 0. 26 0. 05 –0 .1 0 0. 94 0. 89 0. 79 0. 94 0. 80 0. 98 0. 91 1. 00 1. 00 1. 00 sm –0 .5 1 0. 01 0. 22 –0 .6 4 –0 .3 4 –0 .3 0 –0 .3 1 0. 63 0. 93 –0 .6 8– 0. 18 0. 26 0. 05 –0 .1 2 0. 94 0. 90 0. 79 0. 95 0. 81 0. 98 0. 91 1. 00 1. 00 1. 00 1. 00 eu –0 .5 8 –0 .0 6 0. 24 –0 .5 6 –0 .4 1 –0 .4 4 –0 .4 3 0. 66 0. 97 –0 .7 5– 0. 11 0. 19 0. 25 –0 .1 2 0. 97 0. 97 0. 92 0. 90 0. 91 0. 95 0. 98 0. 96 0. 96 0. 96 0. 96 1. 00 g d –0 .5 0 0. 02 0. 22 –0 .6 6 –0 .3 4 –0 .2 7 –0 .2 9 0. 61 0. 92 –0 .6 7– 0. 22 0. 26 0. 04 –0 .1 3 0. 93 0. 88 0. 77 0. 96 0. 78 0. 97 0. 89 0. 99 1. 00 1. 00 1. 00 0. 95 1. 00 d y –0 .5 9 –0 .0 5 0. 35 –0 .7 4 –0 .4 1 –0 .3 7 –0 .4 0 0. 64 0. 94 –0 .7 6– 0. 24 0. 26 –0 .0 1– 0. 22 0. 94 0. 89 0. 76 0. 99 0. 78 0. 96 0. 89 0. 98 0. 98 0. 98 0. 99 0. 95 0. 99 1. 00 h o –0 .6 2 –0 .0 8 0. 41 –0 .7 7 –0 .4 4 –0 .4 0 –0 .4 4 0. 64 0. 94 –0 .7 9– 0. 26 0. 25 –0 .0 3– 0. 26 0. 93 0. 87 0. 74 0. 99 0. 76 0. 95 0. 87 0. 97 0. 97 0. 97 0. 97 0. 93 0. 98 1. 00 1. 00 er –0 .6 4 –0 .1 0 0. 45 –0 .8 0 –0 .4 6 –0 .4 2 –0 .4 7 0. 64 0. 93 –0 .8 1– 0. 28 0. 25 –0 .0 6– 0. 29 0. 92 0. 86 0. 72 1. 00 0. 75 0. 93 0. 86 0. 95 0. 95 0. 95 0. 96 0. 92 0. 96 0. 99 1. 00 1. 00 tm –0 .6 7 –0 .1 2 0. 50 –0 .8 2 –0 .4 9 –0 .4 5 –0 .5 1 0. 64 0. 93 –0 .8 3– 0. 28 0. 25 –0 .0 7– 0. 32 0. 91 0. 86 0. 71 1. 00 0. 74 0. 91 0. 85 0. 93 0. 93 0. 93 0. 94 0. 90 0. 94 0. 98 0. 99 1. 00 1. 00 yb –0 .6 9 –0 .1 4 0. 53 –0 .8 3 –0 .5 0 –0 .4 8 –0 .5 3 0. 64 0. 93 –0 .8 5– 0. 29 0. 24 –0 .0 6– 0. 33 0. 91 0. 86 0. 72 0. 99 0. 75 0. 90 0. 85 0. 92 0. 92 0. 92 0. 93 0. 90 0. 93 0. 98 0. 99 0. 99 1. 00 1. 00 lu –0 .7 1 –0 .1 6 0. 56 –0 .8 4 –0 .5 3 –0 .5 1 –0 .5 6 0. 64 0. 92 –0 .8 7– 0. 30 0. 22 –0 .0 5– 0. 36 0. 90 0. 85 0. 71 0. 98 0. 74 0. 88 0. 84 0. 90 0. 90 0. 90 0. 91 0. 89 0. 91 0. 96 0. 98 0. 99 1. 00 1. 00 1. 00 h f –0 .5 4 –0 .1 2 0. 18 –0 .4 2 –0 .3 7 –0 .4 9 –0 .4 6 0. 69 0. 94 –0 .7 0– 0. 08 0. 19 0. 37 –0 .0 8 0. 88 0. 96 0. 96 0. 77 0. 99 0. 83 0. 98 0. 85 0. 85 0. 84 0. 85 0. 94 0. 82 0. 82 0. 80 0. 79 0. 78 0. 78 0. 78 1. 00 ta –0 .4 8 –0 .0 5 0. 19 –0 .5 4 –0 .3 3 –0 .3 3 –0 .3 5 0. 62 0. 89 –0 .6 4– 0. 08 0. 23 0. 12 –0 .1 3 0. 95 0. 89 0. 80 0. 90 0. 78 1. 00 0. 90 0. 97 0. 98 0. 98 0. 98 0. 95 0. 97 0. 96 0. 94 0. 92 0. 90 0. 88 0. 86 0. 83 1. 00 pb –0 .6 2 –0 .0 9 0. 19 –0 .4 8 –0 .4 4 –0 .5 7 –0 .5 3 0. 81 0. 96 –0 .7 5 0. 04 0. 16 0. 24 –0 .1 5 0. 94 0. 98 0. 91 0. 83 0. 92 0. 90 0. 97 0. 91 0. 91 0. 90 0. 90 0. 96 0. 88 0. 88 0. 86 0. 85 0. 84 0. 84 0. 83 0. 95 0. 90 1. 00 th –0 .6 3 –0 .0 8 0. 24 –0 .5 2 –0 .4 6 –0 .5 5 –0 .5 2 0. 75 0. 98 –0 .7 9– 0. 07 0. 16 0. 29 –0 .1 7 0. 94 0. 99 0. 94 0. 85 0. 96 0. 89 0. 99 0. 91 0. 91 0. 90 0. 91 0. 97 0. 89 0. 89 0. 88 0. 87 0. 86 0. 87 0. 86 0. 98 0. 89 0. 98 1. 00 u –0 .5 7 –0 .0 6   0. 19 –0 .5 2 –0 .4 0 –0 .4 7 –0 .4 4 0. 72 0. 97 –0 .7 4– 0. 15 0. 19 0. 29 –0 .1 4 0. 92 0. 98 0. 92 0. 86 0. 96 0. 89 0. 99 0. 93 0. 92 0. 92 0. 92 0. 98 0. 91 0. 90 0. 89 0. 87 0. 86 0. 86 0. 85 0. 98 0. 89 0. 97 0. 99 1. 00 (c on tin ue d) https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 58 of 139 geusbulletin.org ta bl e 17 (c on tin ue d) p ea rs on p ro du ct m om en t c or re la tio n co effi ci en t ( r) m at ric es fo r m in er al m od es a nd s el ec te d tr ac e el em en ts   su bz on es u za to l zc f o l o px cp x pl m t il ap q o r co sc v cr n i zn rb sr y zr n b ba la ce n d sm eu g d d y h o er tm yb lu h f ta pb th u f 1. 00 o l 0. 27 1. 00 o px 0. 89 0. 23 1. 00 cp x –0 .7 4 –0 .4 7 –0 .7 3 1. 00 pl 0. 57 0. 32 0. 65 –0 .6 9 1. 00 m t –0 .4 8 –0 .3 3 –0 .6 6 0. 34 –0 .6 3 1. 00 il –0 .5 0 –0 .3 8 –0 .6 2 0. 36 –0 .8 2 0. 66 1. 00 ap 0. 29 0. 44 0. 22 –0 .3 3 0. 21 –0 .1 5 –0 .2 2 1. 00 q   o r co –0 .6 4 –0 .0 5 –0 .7 2 0. 39 –0 .7 4 0. 65 0. 80 –0 .1 5 1. 00 sc –0 .7 8 –0 .5 3 –0 .7 6 0. 86 –0 .8 7 0. 52 0. 72 –0 .3 8 0. 61 1. 00 v –0 .2 5 –0 .4 2 –0 .4 7 0. 32 –0 .7 3 0. 68 0. 82 –0 .1 2 0. 60 0. 58 1. 00 cr 0. 50 0. 04 0. 54 –0 .2 7 0. 51 –0 .4 8 –0 .5 8 0. 47 –0 .6 1– 0. 47 –0 .2 7 1. 00 n i 0. 91 0. 27 0. 77 –0 .6 7 0. 36 –0 .3 1 –0 .3 1 0. 29 –0 .3 9– 0. 65 0. 00 0. 49 1. 00 zn –0 .7 7 0. 02 –0 .7 5 0. 40 –0 .5 7 0. 58 0. 61 –0 .2 2 0. 90 0. 54 0. 33 –0 .7 0– 0. 61 1. 00 rb 0. 67 0. 57 0. 71 –0 .6 7 0. 75 –0 .6 5 –0 .7 2 0. 22 –0 .5 7– 0. 83 –0 .7 3 0. 27 0. 51 –0 .4 3 1. 00 sr 0. 23 0. 46 0. 35 –0 .5 2 0. 86 –0 .4 5 –0 .7 4 0. 20 –0 .5 4– 0. 73 –0 .8 1 0. 22 0. 01 –0 .2 3 0. 69 1. 00 y 0. 45 0. 50 0. 56 –0 .2 7 0. 52 –0 .6 6 –0 .7 5 0. 15 –0 .5 2– 0. 56 –0 .7 6 0. 20 0. 26 –0 .3 3 0. 82 0. 52 1. 00 zr 0. 54 0. 37 0. 52 –0 .4 6 0. 25 –0 .3 9 –0 .2 2 0. 06 –0 .1 6– 0. 48 –0 .3 4– 0. 08 0. 41 –0 .1 4 0. 70 0. 19 0. 70 1. 00 n b –0 .3 1 –0 .1 6 –0 .3 3 0. 15 –0 .5 4 0. 30 0. 70 –0 .3 1 0. 61 0. 43 0. 29 –0 .7 2– 0. 28 0. 58 –0 .2 0– 0. 37 –0 .2 2 0. 28 1. 00 ba 0. 28 0. 57 0. 42 –0 .4 8 0. 76 –0 .5 6 –0 .7 4 0. 19 –0 .4 5– 0. 72 –0 .8 5 0. 16 0. 08 –0 .1 5 0. 85 0. 90 0. 78 0. 50 –0 .2 2 1. 00 la 0. 75 0. 56 0. 78 –0 .6 7 0. 74 –0 .6 4 –0 .7 6 0. 28 –0 .6 1– 0. 87 –0 .7 0 0. 31 0. 57 –0 .4 9 0. 95 0. 64 0. 87 0. 75 –0 .2 7 0. 80 1. 00 ce 0. 75 0. 56 0. 78 –0 .6 7 0. 73 –0 .6 5 –0 .7 6 0. 28 –0 .6 2– 0. 86 –0 .7 0 0. 31 0. 57 –0 .5 0 0. 95 0. 63 0. 88 0. 76 –0 .2 7 0. 80 1. 00 1. 00 n d 0. 71 0. 55 0. 75 –0 .6 0 0. 69 –0 .6 6 –0 .7 6 0. 24 –0 .6 0– 0. 81 –0 .7 1 0. 28 0. 52 –0 .4 8 0. 94 0. 60 0. 92 0. 77 –0 .2 5 0. 80 0. 99 0. 99 1. 00 sm 0. 65 0. 55 0. 71 –0 .5 0 0. 65 –0 .6 8 –0 .7 8 0. 23 –0 .6 0– 0. 74 –0 .7 4 0. 28 0. 46 –0 .4 6 0. 92 0. 59 0. 96 0. 75 –0 .2 7 0. 80 0. 97 0. 98 0. 99 1. 00 eu –0 .1 0 0. 48 0. 10 –0 .1 7 0. 52 –0 .4 1 –0 .5 6 0. 02 –0 .2 1– 0. 39 –0 .8 1– 0. 10 –0 .2 9 0. 17 0. 61 0. 80 0. 67 0. 34 –0 .0 4 0. 91 0. 54 0. 54 0. 56 0. 60 1. 00 g d 0. 56 0. 52 0. 62 –0 .4 0 0. 56 –0 .6 3 –0 .7 2 0. 15 –0 .5 3– 0. 65 –0 .7 2 0. 15 0. 36 –0 .3 7 0. 87 0. 54 0. 98 0. 77 –0 .1 7 0. 78 0. 93 0. 93 0. 97 0. 98 0. 63 1. 00 d y 0. 47 0. 49 0. 57 –0 .2 9 0. 52 –0 .6 5 –0 .7 4 0. 14 –0 .5 3– 0. 57 –0 .7 5 0. 17 0. 27 –0 .3 4 0. 83 0. 52 1. 00 0. 72 –0 .2 0 0. 77 0. 88 0. 89 0. 93 0. 97 0. 66 0. 99 1. 00 h o 0. 41 0. 51 0. 51 –0 .2 5 0. 49 –0 .6 3 –0 .7 2 0. 13 –0 .4 8– 0. 54 –0 .7 6 0. 13 0. 21 –0 .2 9 0. 81 0. 53 1. 00 0. 71 –0 .1 8 0. 79 0. 86 0. 87 0. 91 0. 95 0. 70 0. 98 1. 00 1. 00 er 0. 43 0. 50 0. 53 –0 .2 5 0. 50 –0 .6 4 –0 .7 3 0. 14 –0 .4 9– 0. 54 –0 .7 5 0. 15 0. 22 –0 .3 0 0. 81 0. 52 1. 00 0. 71 –0 .2 0 0. 78 0. 86 0. 87 0. 91 0. 95 0. 68 0. 98 1. 00 1. 00 1. 00 tm 0. 38 0. 53 0. 49 –0 .2 2 0. 48 –0 .6 4 –0 .7 3 0. 13 –0 .4 6– 0. 52 –0 .7 6 0. 14 0. 19 –0 .2 8 0. 79 0. 52 0. 99 0. 69 –0 .2 1 0. 78 0. 84 0. 85 0. 89 0. 94 0. 70 0. 97 0. 99 0. 99 0. 99 1. 00 yb 0. 40 0. 50 0. 51 –0 .2 4 0. 48 –0 .6 3 –0 .7 0 0. 13 –0 .4 5– 0. 52 –0 .7 6 0. 13 0. 20 –0 .2 6 0. 81 0. 51 0. 99 0. 72 –0 .1 5 0. 78 0. 85 0. 86 0. 90 0. 94 0. 70 0. 97 0. 99 1. 00 1. 00 0. 99 1. 00 lu 0. 35 0. 49 0. 47 –0 .2 0 0. 45 –0 .6 2 –0 .6 9 0. 11 –0 .4 3– 0. 48 –0 .7 6 0. 09 0. 15 –0 .2 3 0. 78 0. 50 0. 99 0. 71 –0 .1 3 0. 77 0. 83 0. 83 0. 88 0. 93 0. 72 0. 96 0. 99 1. 00 0. 99 0. 99 1. 00 1. 00 h f 0. 56 0. 36 0. 53 –0 .4 5 0. 22 –0 .4 1 –0 .2 0 0. 07 –0 .1 6– 0. 46 –0 .3 1– 0. 08 0. 45 –0 .1 6 0. 70 0. 15 0. 69 0. 99 0. 31 0. 47 0. 75 0. 75 0. 77 0. 75 0. 31 0. 76 0. 72 0. 70 0. 71 0. 68 0. 72 0. 70 1. 00 ta –0 .5 9 –0 .1 0 –0 .5 8 0. 33 –0 .5 8 0. 37 0. 70 –0 .3 1 0. 75 0. 55 0. 25 –0 .7 7– 0. 54 0. 77 –0 .3 1– 0. 30 –0 .2 7 0. 11 0. 93 –0 .1 8– 0. 40 –0 .4 0– 0. 38 –0 .3 7 0. 09 –0 .2 7– 0. 26 –0 .2 2– 0. 24 –0 .2 2– 0. 19 –0 .1 5 0. 12 1. 00 pb 0. 91 0. 13 0. 88 –0 .6 8 0. 52 –0 .4 5 –0 .4 5 0. 19 –0 .7 0– 0. 66 –0 .2 0 0. 47 0. 78 –0 .7 6 0. 55 0. 18 0. 38 0. 44 –0 .2 8 0. 21 0. 64 0. 65 0. 62 0. 56 –0 .1 4 0. 49 0. 41 0. 34 0. 35 0. 30 0. 32 0. 28 0. 46 –0 .5 8 1. 00 th 0. 78 0. 55 0. 78 –0 .6 7 0. 71 –0 .6 2 –0 .7 3 0. 28 –0 .5 8– 0. 86 –0 .6 6 0. 32 0. 62 –0 .5 0 0. 94 0. 60 0. 84 0. 73 –0 .2 4 0. 76 0. 98 0. 98 0. 97 0. 94 0. 49 0. 90 0. 85 0. 82 0. 82 0. 80 0. 81 0. 78 0. 73 –0 .3 9 0. 65 1. 00 u 0. 76 0. 57 0. 77 –0 .6 9 0. 71 –0 .6 1 –0 .7 2 0. 28     –0 .5 7– 0. 87 –0 .6 6 0. 30 0. 59 –0 .4 7 0. 93 0. 61 0. 85 0. 77 –0 .2 3 0. 78 0. 99 0. 99 0. 98 0. 95 0. 52 0. 91 0. 86 0. 84 0. 84 0. 81 0. 83 0. 80 0. 76 –0 .3 7 0. 65 0. 98 1. 00 f: m el t r em ai ni ng . o l: ol iv in e m od e. o px : o rt ho py ro xe ne m od e. c px : c lin op yr ox en e m od e. p l: pl ag io cl as e m od e. m t: m ag ne tit e m od e. il : i lm en ite m od e. a p: a pa tit e m od e. q : q ua rt z m od e. o r: or th oc la se m od e. m in er al m od es a fte r ta bl e 2 an d tr ac e el em en t c on ce nt ra tio ns a fte r t ab le s 14 a nd 1 5. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 59 of 139 geusbulletin.org plagioclase, magnetite, ilmenite and apatite, but positively correlated with quartz and orthoclase. this suggests that the increase in the excluded elements in the upper part of the intrusion may be associated with an evolved melt component and not with the predominantly silicate and oxide building blocks of the gabbros. the lower part of the ls shows mixed results, probably due to the complexity of this gabbro sequence that consists of several zones and subzones. again, the excluded elements show positive interdependency, except for a negative correlation between both ta and nb and the rest of the group. these two elements are further positively correlated to ilmenite, magnetite and clinopyroxene (table 17). most included elements are negatively correlated with the excluded elements, except cr and ni that are weakly positively correlated. furthermore, the excluded elements are once again positively correlated with f, meaning that they are increasing upward in the stratigraphy due to differentiation. this indicates a decoupled evolution of the two groups of included and excluded trace elements controlled by differences in their bulk partition coefficients (see discussion in section 5.11). there are no strong positive correlations in the lz to uza between trace element concentrations and mineral modes, except for sr and ba in plagioclase and sc in clinopyroxene. remarkably, olivine and apatite do not show strong positive correlations with any trace elements (included or excluded). for apatite, this probably reflects the low modal content of this mineral, while for olivine, most included elements are variably partitioned into all mafic and feti oxide minerals. it is furthermore unexpected that the included transition elements do not appear to act as a coherent group of elements with strong positive intra-correlations. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 60 of 139 geusbulletin.org 5 discussion 5.1 liquidus temperatures there have been several attempts to estimate the liquidus temperatures of the skaergaard intrusion. experimental methods have constrained the upper bracket by determining the liquidus conditions for several potential initial melt compositions. wager (1960, 1961) identified a sample interpreted as a chilled margin of the intrusion and, thus, representative of the skaergaard magma. tilley et al. (1963) and biggar (1974) melted this sample at atmospheric pressure using a nno oxygen buffer and determined the liquidus at 1230 ± 16°c, co-saturated with olivine and plagioclase. subsequent petrographic studies (mcbirney 1975, 1996) identified a more suitable chilled margin sample that was found to have plagioclase on the liquidus at 1175 ± 5°c for the magnetite-wüstite (mw) oxygen buffer (hoover 1989a). snyder et al. (1993) melted the same chilled margin within the stability field of feti oxides (1091–1062°c). toplis & carroll (1995) and thy et al. (2006) used chilled margins of dykes from a swarm believed to be associated with the skaergaard intrusion (fg-1 dyke swarm; brooks & nielsen 1978, 1990). toplis & carroll (1995) investigated an oxide mixture approximating the c-dyke composition (lacking mno and p2o5) and found that plagioclase and olivine appeared on the liquidus at 1162 ± 4°c. thy et al. (2006) bracketed the liquidus for the actual chilled margin of the same dyke at 1171 ± 3°c and found plagioclase at the liquidus, followed by near co-saturation of olivine and clinopyroxene at 1149 ± 10°c. mcbirney & naslund (1990) melted a series of gabbros considered representative of the ls between 1150°c (lza) and 1002°c (uzc) at various values of t and fo2 estimated by morse et al. (1980) and williams (1971), and thus do not provide independent temperature estimates. the lower temperature bracket for the uzc was constrained by tilley et al. (1963) using direct crystallisation experiments on a fayalitic diorite from the middle of uzc (wager & brown 1967). they found plagioclase on the liquidus at 1035°c, with bustamite and ferrohedenbergite subsequently appearing at 1010°c and inferred the liquidus temperature to have been above the 970°c inversion of ferrobustamite to ferrohedenbergite (yoder et al. 1963) and below the 1035°c plagioclase liquidus of the ferrodiorite. lindsley et al. (1969) used the ferrobustamite-ferrohedenbergite and the tridymite-quartz inversions to further constrain crystallisation temperatures of 980–950°c for the sh at a pressure of 600 ± 100 bars. these experimental constraints suggest that the liquidus temperatures for skaergaard magma decreased by 206–265°c from its initial emplacement to final solidification represented by the sh, depending on the choice of initial magma composition. forward modelling results of toplis & carroll (1996) and thy et al. (2006, 2008, 2009a) largely recorded the abovequoted experimental temperatures for the respective initial melt compositions, but were unable to constrain the terminal liquidus temperature. ariskin (2002, 2003) used a crystallisation model based on existing experimental low-pressure information for basaltic systems to constrain the temperature variation in the skaergaard intrusion. his conclusion was that the initial magma filling the chamber was olivine and plagioclase-phyric, which in the lza had equilibrated at 1145°c. several other attempts to constrain the temperature bracket for the skaergaard intrusion have with various successes been attempted. hess (1941) estimated a bracket of 185°c, constrained at high temperature (c. 1140°c) by the orthoto clinopyroxene inversion and at low temperature (955°c) by the hedenbergite-ferrobustamite inversion. williams (1971) used olivine-pyroxene-oxide equilibria of cumulates as well as lindsley et al.’s (1969) terminal liquidus temperature to suggest a bracket of c. 400°c (1300–900°c), apparently including the large hz of wager & deer (1939). the temperature bracket of 370°c (1320–950°c) estimated by kudo & weill (1970) was based on experimental plagioclase-liquid thermometry and again extrapolated to include the large hz originally suggested by wager & deer (1939). by restricting the bracket of kudo & weill (1970) to the exposed cumulate column, a range of 270°c (1220– 950°c) is obtained. this estimate is rather close to that of morse et al. (1980) suggesting 300°c (1250–950°c), based on the plagioclase temperatures determined of kudo & weill (1970) and a terminal liquidus temperature of 950°c (lindsley et al. 1969). further attempts to constrain liquidus temperatures for skaergaard have relied on singleand two-mineral geothermometers (fig. 39). morse (2008a) used a 5 kbar study of the kiglapait intrusion to model the plagioclase liquidus of the skaergaard intrusion using the plagioclase compositions of toplis et al. (2008). this suggested lza/ lzb liquidus temperature of 1173°c or extrapolated to the hz of 1211°c (both temperatures adjusted to atmospheric pressure). thy et al. (2009b) proposed an empirical single-mineral plagioclase thermometer based on a large data set of melting experiments for north atlantic basalts with coexisting plagioclase and melt at atmospheric pressure. they obtained a liquidus temperature of 1155°c for the hz, falling to 1140°c in the lzb, 1050°c in the uzc and finally 1000–1025°c in the sh. these terminal temperatures were similar to those predicted by https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 61 of 139 geusbulletin.org plagioclase augite-olivine pigeonite-augite magnetite-ilmenite olivine augite 600 650 700 750 800 850 900 950 1000 1050 1100 1150 0 200 400 600 800 1000 1200 1400 1600 1800 2000 2200 stratigraphic height (m) t (º c ) lzb uza lzc mz pl mt-il pi-aug aug-ol aug ol ab fig. 39 summary of calculated equilibrium temperatures (t; °c) using singleand two-phase thermometers as a function of stratigraphic height (m) in the ls. plagioclase temperature (t = 3.6 × an + 899) is obtained using the formula of thy et al. (2009b), with a 1σ sd based on multiple analyses of plagioclase grains. the extension to 920°c for albite (ab) is taken to represent the crystallisation of the interstitial granophyre. olivine and augite temperatures are calculated using linear regressions of the same experimental dataset of thy et al. (2006), respectively, t = 3.4 × fo + 910 (σr2 = 0.943) and t = 5.4% × mg# + 750 (σr2 = 0.872), assuming all iron is fe2+. the coexisting augite-olivine temperature is obtained using the formula of loucks (1996). the coexisting pigeonite-augite temperature is after putirka (2008; their eq. 37), recommended for pyroxenes with a magnesium number (mg#) < 0.75. equilibrium temperatures for coexisting magnetite-ilmenite are from table 9, calculated using the formula of ghiorso & evans (2008). the lzc–mz interval discussed in the text is shown in the centre of the diagram. the mineral compositions used for the calculations of crystallisation temperatures are from tables 3–5, 8 and 9. pl: plagioclase. ol: olivine. aug: augite. pi: pigeonite. mt: magnetite. il: ilmenite. morse (2008a), liquidus temperatures from the thy et al. (2009b) plagioclase geothermometer for the hz and much of the ls and is >50°c below those estimated using the morse (2008a) geothermometer (fig. 39). because of sluggish reaction kinetics and difficulties in attaining and demonstrating equilibrium of plagioclase, specifically in dry melting experiments and at low temperatures, attention is required when products are evaluated and when experimental plagioclase compositions are used in petrogenetic modelling, as in this study (johannes & koepke 2001; morse 2010). attention to this problem was highlighted by morse (2010) and subsequently addressed by thy et al. (2013) for the skaergaard dyke experiments used here. the study by thy et al. (2013) on the effects of non-equilibrium or unreactive plagioclase grains in melting experiments provided tighter bounds on experimental plagioclase composition and documented composition-dependent partitioning of na and ca between plagioclase and melt. application of the results to model the skaergaard, however, requires only minor adjustments to the previously proposed liquidus temperatures and liquid line of descent modelling that does not significantly change the temperature constraints. liquidus temperatures from thy et al. (2009b) are preferred here because they conform to the existing melting experiment on probable initial melt compositions as outlined above (1162–1175°c). similar single-mineral thermometers formulated for olivine and clinopyroxene using the experimental results for the skaergaard dykes of thy et al. (2006) obtain temperatures for the base of lza of 1117°c and 1136°c, respectively, comparable to the 1137°c plagioclase liquidus temperature determined by thy et al. (2009b; see caption to fig. 39). a single-mineral clinopyroxene thermometer reformulated by putirka (2008, their eq. 32d) reproduces the expected temperatures for the two end members from plagioclase compositions reasonably well, but fails to give meaningful results for the intermediate pyroxenes. the reason for this may be related to the unusual low-octahedral occupancy of al for the skaergaard pyroxenes. the temperatures obtained using the two-mineral clinopyroxene-olivine thermometer formulated by loucks (1996) give comparable results to the single-mineral temperatures, while the pigeonite-augite thermometer reformulated by putirka (2008) deviates significantly (fig. 39). the results for the various temperature estimates are compared in fig. 39 as a function of stratigraphic height. the plagioclase estimates are systematically higher throughout the ls than those based on the mafic minerals. compared to plagioclase, the mafic minerals (olivine and pyroxenes) record lower equilibration temperatures in lza and lzb. in addition, the equilibration temperatures in the uz markedly decrease upward into uzc where the estimates of clinopyroxene temperature drops. the two-pyroxene thermometer (pi-aug; pigeonite-augite) records equilibration temperatures of 1100–1075°c in the lz and into mz. all the mafic mineral thermometers (olivine (ol), augite (au) and au-ol) show corresponding and decreasing results upward in the uz (fig. 39). if the plagioclase thermometer is applied to the albitic plagioclase (ab) in the granophyre (figs 13, 14), a terminal temperature is obtained down to c. 920°c, https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 62 of 139 geusbulletin.org largely conforming to that obtained for olivine and based on the same melting experiments. the apatite saturation temperature of 1060°c, estimated using the equation of tollari et al. (2006; see section 5.7.8), also corresponds well to the plagioclase solidus at the uza– uzb boundary. the estimated liquidus temperature converges to the solidus temperature in the lzc–mz interval, recording little, if any, subliquidus equilibration. this contrasts to the lza and lzb below and the uz above (fig. 39). this is probably related to the low-trapped melt content in the lzc–mz interval as discussed in section 5.7.6. the ‘dome’ in the near solidus estimates is also reflected in the mostly subsolidus equilibration and oxidation temperatures for coexisting magnetite-ilmenite pairs. the feti oxides, magnetite and ilmenite, first appear in lzc as early primocrysts at c. 1100°c. temperatures obtained for coexisting ilmenite and magnetite (ghiorso & evans 2008) are at least 200–300°c below the experimental crystallisation temperatures obtained by thy & lofgren (1994) for ferrobasaltic melts, and are well into the subsolidus range several hundreds of degrees below magmatic conditions (fig. 39). several observations from fig. 39 are not easily explained. low temperatures in the lz and uz for the mafic minerals, compared to the higher values for plagioclase, may be related to the presence of high amounts of trapped melt (see sections 5.11.4 and 5.12.6). this is the “trapped melt effect” suggested by barnes (1986) to be the result of modification of mafic mineral compositions from reactions with or crystallisation from trapped interstitial melt. the mafic minerals, like olivine, are far more susceptible to diffusional adjustment than plagioclase, which may not be affected within the likely time scales of processes acting in a magma chamber, such as the cooling in the skaergaard intrusion (morse 1984; grove et al. 1984). this is certainly the case for the pyroxenes where solidus and subsolidus re-equilibration and exsolution (see sections 4.3.3–4.3.5) may have reset mineral compositions and thus the obtained temperatures. likewise, the effects of the onset of liquid immiscibility in the upper part of the intrusion (uzb; see section 5.12) may also have affected the pyroxene equilibria (bustamite and ferrohedenbergite in uzc) and the estimated crystallisation temperatures. 5.2 oxygen fugacity (fo2) the temperature and fo2 constraints during crystallisation and solidification of the skaergaard intrusion are, and have been, a subject of some contention. the most common approach to constraining fo2 for evolved gabbroic rocks is the magnetite-ilmenite oxy-geothermometer (buddington & lindsley 1964). we employ this approach using the calibration of ghiorso & evans (2008) with the understanding that fo2 estimates reflect conditions at magnetite-ilmenite closure temperatures and not necessarily the liquidus. these results are summarised in table 9 and illustrated in fig. 40. the highest temperature determinations cluster around the nno oxygen buffer at 800–850°c (fig. 41). below about 700°c, the variation clusters around the fmq oxygen buffer to very low temperatures of 400°c or below (table 9; figs 40, 41). these temperature estimates are well below any reasonable estimates for the liquidus appearance of coexisting ilmenite and magnetite in the skaergaard intrusion (thy & lofgren 1994; toplis & carroll 1996; thy et al. 2006) and those implied by the thermometry of the coexisting silicates (fig. 39). similar results were obtained by bollingberg (1995) using an extensive, but different sampling of gabbros through the ls. low temperature determinations for coexisting feti oxides have been related to subsolidus oxidation of ulvöspinel to ilmenite and its granule exsolution coarsening as secondary ilmenite (vincent 1960; buddington & lindsley 1964). reconstruction of the primary equilibrium oxide compositions has been an integral process for estimates –26 –24 –22 –20 –18 –16 –14 –12 –10 –8 400 500 600 700 800 900 1000 1100 1200 t (ºc) fe-enrichment si-enrichment m w n n o fm q o xy ge n fu ga ci ty (l og fo 2) morse et al. (1980) sh ariskin (2003) buddington & lindsley (1964) williams (1971) morse et al. (1980) kersting et al. (1989) sato & valenza (1980) lindsley et al. (1969): sh this study fig. 40 summary of subsolidus oxygen fugacity (log fo2) vs. temperature (t) for coexisting magnetite and ilmenite using the thermobarometer of ghiorso & evans (2008). the results of this study (black dots; table 9) are compared to similar results of buddington & lindsley (1964; open inverted triangles) and to selected oxygen buffers (nno, fmq and mw). the subsolidus results are compared to various calculated and measured estimates (see section 5.2 for details). the large blue arrow is a visual interpretation of the overall observed variation. the morse et al. (1980) trend line is based on a combination of mineral observations and thermodynamic calculations. the si-enrichment trend line (small black dots) is from supplementary table s1, and the fe-enrichment trend line (small grey dots) is from supplementary table s2. note that the results of williams (1971) extend to 1300°c but are truncated here. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 63 of 139 geusbulletin.org of temperature and fo2. buddington & lindsley (1964) used the bulk compositions of separated magnetite and ilmenite from the ls (vincent & phillips 1954) to approximate the primary compositions. the temperatures obtained by buddington & lindsley (1964), varying from 900°c to 775°c and clustering around the fmq oxygen buffer (fig. 40), were nevertheless interpreted to reflect subsolidus oxidation with an intergranular fluid action and not late-stage magmatic melts. further reconstruction was thus required to identify the liquidus conditions. this was attempted by buddington & lindsley (1964) and lindsley et al. (1969) by incorporating granule “oxy-exsolved” ilmenite (vincent & phillips 1954; wright 1961) back into the observed magnetite. the results suggested temperatures from 1150°c to 1025°c for the lzc and 980–950°c for the sh with a corresponding variation in fo2 from the fmq to the mw oxygen buffers. few, if any, of these reconstructed compositions provide information on the primary liquidus conditions. the approach taken in this study has been to analyse the oxide phases with a broad beam or scanning electron beam. this approach does not improve the results, but it allows us to explore the stratigraphic sequence outside the main primary precipitation of ilmenite and magnetite in the lzc and mz (fig. 40). the fo2 is recorded by both the coexisting oxides and in some coexisting silicates (frost et al. 1988; lindsley et  al. 1990; andersen et al. 1993), allowing the uncertainty associated with the coexisting feti oxides to be reduced. this approach has been used to further constrain estimates of fo2 for the skaergaard intrusion based on the mineralogical data of wager & brown (1967), feti oxides and silicate equilibria, and independent evaluations of temperature (e.g. kudo & weil 1970). lindsley et al. (1969) calculated the fo2 –t conditions for the sh using ilmenite-magnetite and 2200 2000 1800 1600 1400 1200 1000 800 600 400 200 0 –2 –1 2 30 1–4 –3900 1000 1100 1200400 500 600 700 800 ~ m w fm q n n o t (°c) delta log fo2 (δfmq) tplag si fe st ra ti gr ap hi c he ig ht (m ) lza lzb lzc mz uza uzb uzc fig. 41 subsolidus temperature (t, °c) and oxygen fugacity (fo2) as a function of stratigraphic height (m). fo2 is on a log basis normalised to fmq (δfmq) and calculated from the analysed coexisting magnetite and ilmenite. tplag: liquidus t, calculated based on plagioclase anorthite content. ~mw, fmq and nno: oxygen buffers, given for comparison. the mw buffer is schematic and does not show the large fluctuation on the log curve attributable to the tplag variation. si: si enrichment of the liquid line of descent (table 21). fe: fe enrichment of the liquid line of descent (table 22). blue horizontal dashed lines: divisions of the ls. abbreviations in fig. 2. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 64 of 139 geusbulletin.org fayalite-magnetite-quartz equilibria. williams (1971) and morse et al. (1980) obtained results largely consistent with the reconstructed magnetite-ilmenite thermometry by buddington & lindsley (1964). the results by morse et al. (1980) provided internally consistent results that fit the terminal conditions of lindsley et al. (1969). these latter results have provided the benchmark fo2 –t variation for the skaergaard intrusion used by most workers as shown in fig. 40, from an initial (lzc) fo2 near the fmq oxygen buffer, followed by more reducing conditions reaching the mw oxygen buffer at the sh. direct measurements of the intrinsic fo2 of skaergaard gabbros and mineral separates were made by sato & valenza (1980) and kersting et al. (1989) using the oxygen-specific, solid-electrolyte cell method. this method provides linear relationships between the experimental measurements for temperature and fo2 and thus still requires an assumption or independent knowledge of the target temperature. sato & valenza (1980) examined a suite of magnetically separated feti oxides representing the zones and subzones of the ls (mcbirney 1989a). their results plot at reduced conditions between the mw and iron-wüstite (iw) oxygen buffers and temperatures of 1130–1000°c (fig. 40). they attribute this to the presence of graphite during crystallisation and the fo2 control of the c-(co+co2) equilibrium, which for an appropriate pressure could explain the observed reduced fo2 (sato & valenza 1980). in a similar study, kersting et al. (1989) used magnetically separated feti oxide mineral concentrates of three gabbros from the collection of wager & brown (1967). the result was consistently at about 0.5 log units above the nno oxygen buffer for oxidising and reducing temperature directions and after argon purging. furthermore, when they examined the final products using feti oxide compositions and buddington & lindsley’s (1964) thermobarometer, they obtained fo2–t relations consistent with their experimental end conditions about one log unit below the electrochemical results between the nno and fmq oxygen buffers. these latter results were their preferred values and are also those shown in fig. 40. the large differences between the two sets of intrinsic fo2 measurements have been attributed to carbon from the original minerals or carbon introduced during laboratory handling and procedures (kersting et al. 1989); however, the issue has never been convincingly settled. melting experiments used to support modelling of the skaergaard liquid lines of decent at atmospheric pressure have been conducted in equilibrium with an fo2 between the iw and nno oxygen buffers and at temperatures from liquidus to near solidus conditions (biggar 1974; hoover 1989a; snyder et al. 1993; toplis & carroll 1995; thy et al. 2006). the initial crystallisation conditions for the skaergaard intrusion used for these various experiments varied from 1170°c to –9.5 log fo2 (hoover 1989a), 1180°c to –11.6 log fo2 (snyder et al. 1993) and 1160°c to –9.4 log fo2 (toplis & carroll 1995, 1996). these estimates range from below the fmq to well below the mw oxygen buffers and are assumptions often based on morse et al. (1980) and thus do not represent independent values. forward modelling using a preferred initial composition and fo2 can nevertheless be used to explore the trajectories for temperature and fo2 of the liquid line of descent assuming various crystallisation models. toplis & carroll (1996) modelled the liquid lines of decent for equilibrium and fractional crystallisation for systems open or closed to the exchange of oxygen with the host gabbros of the intrusion. in systems controlled by fractional crystallisation and open to exchange of oxygen, the modelled fo2 varied parallel to an oxygen buffer defined by the initial conditions, as observed in oceanic volcanic centres. in systems closed to oxygen exchange, as in plutonic centres like the skaergaard, the modelled fo2 varied depending on the type of crystallising minerals and their proportions. the modelling that best fitted the observed cumulus sequence of the intrusion was inferred to be fractional crystallisation under conditions closed with respect to the exchange of oxygen with the surroundings (toplis & carroll 1996; thy et al. 2006). thy et al. (2006, 2009a) explored the effects of the modes of magnetite and ilmenite. they pointed out that the models that assumed an extrapolation of the experimental modes of feti oxides would result in dramatically different trajectories of fo2 –t than those using the observed mode of the skaergaard cumulates. the two contrasting trajectories are shown in fig. 40 marked as either ‘si enrichment’ (experimental mode; toplis & carroll 1996) or ‘fe enrichment’ (observed modes; thy et al. 2009a) controlled by the amounts of feti oxides used in the calculations. both paths are generally consistent with the observed mineral-zone layering, although they indicate dramatically different end-stage melt compositions (see section 5.4). further, fig. 41 illustrates the two trajectories expressed as fo2 on a log basis normalised to fmq (δfmq) with stratigraphic height. the si enrichment trend follows the results obtained for the coexisting ilmenite and magnetite, while the fe enrichment trend is discordant to the observed variation above mz. 5.3 lithostatic pressure the skaergaard magma chamber intruded into the unconformity between the precambrian basement and the overlying tertiary basalts during the peak of continental rifting and build-up of the plateau lava shield (irvine 1991; nielsen 2004). the intrusion is interpreted to have intruded around 56–55 ma, or slightly earlier, https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 65 of 139 geusbulletin.org after eruption of about half of the plateau lavas, and to have cooled relatively quickly until the overburden thickness had reached its maximum extent 5–6 km (hirschmann et al. 1997; barfod et al. 2003; tegner et al. 2008; brooks 2011). holness et al. (2015) suggested that the intrusion originated as a hz sill composed of several crystal-laden influxes that quickly ballooned to its present volume in response to tectonic expansion. the pressure at the level of the sh was estimated by lindsley et al. (1969) at 60 ± 10 mpa using the hedenbergite-bustamite and tridymite-quartz inversions. by assuming a fast inflation of the chamber (holness et al. 2015) and consequently a constant overburden, the pressure at the level of the lza can be estimated at 125– 130 mpa, using a standard lithospheric gradient of 30 mpa km–1 (morse 2008a; thy et al. 2009b). the effects of such modest anhydrous pressures in the range to 130 mpa will have small or undetectable effects on phase compositions and equilibria, mainly as a modest expansion of the calcic pyroxene field at the expense of plagioclase (yang et al. 1996; whitaker et al. 2007; botcharnikov et al. 2008). this observation justifies the use of one-atmosphere phase equilibria as a reasonable proxy for understanding the main features of the evolution the skaergaard intrusion (toplis & carroll 1995; thy et al. 2006). larsen & tegner (2006) estimated pressure by using the intercepts between isochoric fluid inclusion paths for granophyre with the minimum melt solidus for granitic compositions. their results suggested granophyre formation from 180 ± 50 to 330 ± 130 mpa and corresponding temperatures of 690–660°c. they attributed this upward increase in pressure to progressive burial by outpouring of a 5–6 km thick pile of flood basalts during the final stages of cooling following the emplacement of the intrusion as a sub-volcanic chamber. 5.4 end-stage melt compositions constraining the end-stage melt composition is an important prerequisite for understanding the dominant differentiation processes in basaltic magma systems as well as for the foundation and confirmation of liquid line of descent modelling. one line of evidence that may have bearings on the end-melt composition of the skaergaard intrusion is the occurrences of potential late-stage segregated differentiation products (gabbroic to melanocratic or leucocratic granophyres). in the skaergaard intrusion, podiform segregation bodies ranging from gabbroic to granophyric (or rhyolitic) in composition are regarded as contemporaneous melt and their evolution believed to approximate to the possible hydrated liquid line of descent of the intrusion (larsen & brooks 1994). these granophyric compositions show similarities to observed or theoretically predicted end-stage melts from extreme differentiated basalts or melts formed from initial partial melting of basaltic rocks (bowen 1928; tuttle & bowen 1958; carmichael 1964; presnall & bateman 1973; beard & lofgren 1989; thy et al. 1990; ghiorso & carmichael 1995; villiger et al. 2004). the granophyric compositions, in addition to showing similarities to predicted end-stage fractional crystallisation melts, also show similarities to the conjugate silicic melt produced as a result of liquid immiscibility in basaltic systems and the skaergaard magma (mcbirney & nakamura 1974; mcbirney 1975; dixon & rutherford 1979; philpotts 1979, 1982; jakobsen et al. 2005, 2011; charlier & grove 2012; charlier et al. 2013; honour et al. 2019a). although it may not be a simple exercise to distinguish between the two types of possible final silicic melts (fractional crystallisation or liquid immiscibility), the presence of a coexisting iron-rich melt or rock, identified either microscopically or macroscopically, is conclusive evidence for action of liquid immiscible melts; although not necessarily of the magnitude of such an effect. the abundant interstitial granophyric patches in ferrodiorites of the uzc are good candidates for potential end-stage melt compositions originated from liquid immiscibility (figs 13, 14). salmonsen & tegner (2013) reported a granophyre representative for the sh, supporting the contention that granophyre is a final, or one of the final, differentiation products, of the intrusion. granophyre further commonly appears as veins and dyke and sill-like intrusions throughout the ls, particularly in the uz and the adjacent ubs, and is commonly viewed as a late-stage, sometimes water-saturated, segregation or columnar migration from the gabbro mush during compaction or other stress-induced movements (wager & deer 1939; wager & brown 1967; mcbirney 1989a; larsen et al. 1992; larsen & brooks 1994; larsen & tegner 2006; jakobsen et al. 2011). because of their appearance, granophyric veins and inclusions were not directly sampled and analysed in this study. although some uz ferrodoritic samples evidently contain a high amount of interstitial felsic granophyric components localised in the uppermost part of uzc (figs 13, 14), where quartz reaches c. 10% by weight (figs 3, 15) or equivalently c. 20% calculated as granophyre (see sections 5.11.4 and 5.12.6 on trapped melt content). analyses of granophyres in the ls are provided by wager & deer (1939) and mcbirney (1989a). in addition to the interstitial granophyre in the gabbros, intrusive and segregated granophyres also exist both in the skaergaard intrusion and the surrounding host volcanic rocks. these have been referred to as transgressive granophyre dykes (wager & brown 1967; bird et al. 1986; hirschmann 1992) or as major sills, most noticeably the tinden and sydtoppen granophyres (wager & brown 1967). https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 66 of 139 geusbulletin.org despite its stratigraphic setting and petrographic integration into the gabbroic or dioritic rocks of the intrusion, granophyre is not always considered the final fractionation product of the skaergaard magma. the transgressive granophyre dykes are, in particular, interpreted as derived from a skaergaard-like basaltic magma that incorporated fused archean crust (taylor & forester 1979; hirschmann 1992). similarly, naslund (1989) described from the basaltic basistoppen sill, which intruded into the skaergaard, terminal granophyric differentiation products. these were interpreted as resulting from a combination of fractionation of a skaergaard-like magma and assimilation of melted gneissic host rock. wager & deer (1939) further described several granophyric inclusions in the ls and interpreted them as gneissic inclusions. on the other hand, they interpreted the tinden granophyre and intermediate hedenbergite granophyre of the uzc and sh as representing the final differentiation products. the assimilation of small amounts of gneiss-derived melts into the evolving skaergaard magma is well constrained by studies of stable and radiogenic isotopes (hamilton 1963; leeman & dasch 1978; taylor & forester 1979; norton et al. 1984; stewart & depaolo 1990; mcbirney & creaser 2003; hagen-peter et al. 2019; cho et al. 2022). more recently, bindeman et al. (2008) and wotzlaw et al. (2012) examined ferrodiorites around the sh and concluded that the final magma near the sh became locally reduced in δ18o due to an influx of meteoric water derived from blocks of altered ubs assimilated during the time of mz formation. these authors also suggested that the still hot rocks around the sh might have been reheated and partially melted by the emplacement of the thick basistoppen sill only a few hundred metres above. a comprehensive study of the granophyres of the intrusion has, however, not been conducted as part of the present study. nevertheless, the granophyres in the ls range from felsic granophyre to melanocratic granophyre (or melanogranophyre). the melanogranophyres have gained a special status in recent studies having been interpreted as in part recording liquid immiscibility (mcbirney & nakamura 1974; mcbirney 1975, 1989a, 1995, 1996; jakobsen 2007; jakobsen et al. 2011). however, there is no petrographic or structural field evidence of their origin or their relationships to leucocratic granophyres. table 18 summarises existing information available for the compositions of major and trace elements of skaergaard-related granophyres including granophyre (mcbirney 1989a; larsen & brooks 1994; salmonsen & tegner 2013), melanogranophyre (mcbirney 1989a), transgressive granophyre (hirschmann 1992) and the basistoppen sill (naslund 1989). the granophyric rocks contain variable additions of the coexisting primocrysts, reflected by elevated tio2, feo, mgo and cao particularly for the melanogranophyres (table 18). the trace elements, however, have received less attention as summarised in table 18. there is nevertheless a reasonable correspondence between the trace elements, often within an order of magnitude, and the felsic granophyres and the melanogranophyres. typical included and excluded trace elements ratios approaching the sh, like ni/rb, range from 0.06 to 0.15 and 0.14 to 0.22 for v/rb, although melanogranophyres reach 0.76 in the latter, probably due to inclusion of accidental coexisting primocrysts or from reactions between gabbro and granophyre melts forming hybrid material (wager & deer 1939, p. 189). it is thus suggested that granophyre is the only identifiable late-stage silicic component in the skaergaard. wager & deer (1939, p. 112–117) suggested that the final 200 m of rocks of the sh represented a sheet of melt from which initially fayalitic olivine settled to the base and andesitic plagioclase floated to the top. the final solidification of this sheet occurred without fractionation ending up as a fayalite-hedenbergite granophyre, a term that they applied to the major part of the uzc. the central fayalitic-hedenbergitic granophyre of wager & deer (1939) and wager & brown (1967, p. 95) can be “[..] regarded as a mildly intrusive, penecontemporaneous sheet.” they observed that “[..] there still remains a thin layer of rock, of distinctive composition, which is believed to have formed from residual liquid sandwiched between the upper border group and the ls, and which cannot satisfactorily be classified with either of the adjacent groups.” the primocryst phases in this thin layer were described as andesine plagioclase (an~30), fayalitic olivine and hedenbergitic pyroxene. thus, wager & brown (1967, p. 95) argued for an ironrich skaergaard end product that they referred to as ferrodiorite containing zoned plagioclase, brown ferrohedenbergite, fayalite and iron ore with quartz and orthoclase in abundant (5%) silicic mesostasis areas (e.g. in fig. 14). wager & brown (1967) further interpreted melanogranophyres, granophyres, and some acid granophyres as melts that were segregated from the extreme ferrodioritic fractionation product of the sh or that the mesostasis areas of the ferrodiorites in their view represented the final differential product. mcbirney (1975, 1989a) suggested that only melanogranophyres and what he referred to as segregation granophyric veins and sills, both types occurring in the uzc, were true silicic differentiation products from the skaergaard magma that resulted from liquid immiscibility followed by buoyant separation and accumulation. despite these early observations, in subsequent papers mcbirney did not argue for liquid immiscibility during the evolution of the skaergaard magma (mcbirney 1995, 1996). https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 67 of 139 geusbulletin.org observations by wager & brown (1967) and mcbirney (1989a), as well as those presented here, suggest that the final melt composition at the sh level was in equilibrium with oligoclase (an~30), fayalitic olivine and hedenbergite. the interstitial granophyric melt crystallised to form a fine-grained assemblage of quartz, orthoclase and albitic plagioclase (fig. 14). although it is possible that the interstitial granophyre could be in table 18 summary of average skaergaard granophyric compositions compared to modelled compositions oxide (wt%) granophyre transgressive granophyre basis toppen melanogranophyre sandwich horizon (sh) uzb–uzc boundary ave. granite n ave. sd n ave. sd ave. (n = 2) n ave. sd gabbros analysed melt (si) modelled melt (fe) modelled gabbros analysed melt (si) modelled melt (fe) modelled sio2 19 74.30 3.07 15 69.25 3.57 67.55 7 57.35 3.06 50.17 74.43 55.89 45.64 72.68 55.57 72.04 tio2 19 0.66 0.32 15 0.46 0.29 0.67 7 2.07 0.38 1.87 0.10 0.92 2.54 0.10 1.03 0.30 al2o3 19 12.11 0.68 15 14.23 0.80 11.58 7 11.21 0.98 9.40 10.01 5.04 8.98 10.06 5.57 14.42 fe2o3 10 2.59 0.62 15 1.16 0.42 2.81 4 3.38 1.91 6.28 0.12 8.73 2.68 0.30 8.20 1.22 feo 10 3.25 2.11 15 2.68 1.85 5.30 4 11.85 0.75 18.45 9.61 21.20 23.96 10.27 21.01 1.68 fe2o3* 19 5.82 2.75 15 4.14 2.39 8.69 7 15.05 2.91 26.78 10.81 32.30 29.31 11.71 31.55 1.68 mno 19 0.09 0.04 15 0.07 0.03 0.16 7 0.23 0.05 0.19 0.29 0.30 0.74 0.29 0.31 0.05 mgo 19 0.32 0.30 15 0.53 0.41 2.75 7 1.41 0.48 0.03 0.10 0.40 0.68 0.10 0.40 0.71 cao 19 1.55 0.84 15 1.87 1.11 4.36 7 4.82 0.96 7.89 1.85 2.00 9.65 2.33 2.00 1.82 na2o 19 4.38 0.34 15 4.63 0.27 1.87 7 3.84 0.47 3.25 4.08 2.27 2.68 4.03 2.40 3.69 k2o 19 2.97 0.78 15 3.93 0.74 2.87 7 1.76 0.20 0.80 3.28 3.91 0.32 3.12 3.65 4.12 p2o5 19 0.12 0.12 15 0.20 0.30 0.09 7 0.55 0.23 0.36 0.10 0.10 0.71 0.10 0.10 0.12 total 102.35 99.00 97.13 98.47 98.70 103.98 100.77 98.57 103.38 100.24 100.17 trace elements (ppm) inverted inverted inverted inverted co 13 9 6 12 6 6 2 7 26 4 3 1 3 37 20 45 sc 12 6 4 8 44 2 3 37 4 8 v 12 10 5 8 19 10 9 3 30 52 24 10 26 6 2 5 cr 7 7 2 11 9 4 23 5 7 2 22 0.2 1 10 0.1 0.3 ni 9 10 6 15 5 5 9 5 4 3 4 0.0 0.1 1 0.0 0.0 zn 19 79 48 15 71 25 253 7 161 70 259 250 630 153 183 417 rb 22 69 20 15 87 23 65 7 40 7 16 76 147 5 39 83 sr 23 142 46 15 156 92 173 7 240 35 570 296 354 298 255 359 y 15 152 34 13 39 63 145 3 128 20 117 79 142 62 40 83 zr 19 971 308 12 311 95 1275 7 568 175 398 140 361 77 44 179 nb 6 93 23 12 30 9 90 52 20 51 11 6 14 cs 12 0.8 0.3 ba 19 665 196 15 807 208 805 7 473 109 349 462 435 87 1903 1482 la 6 138 51 12 64 9 93 1 51 35 60 121 13 27 61 ce 6 203 56 12 134 16 201 98 125 245 35 57 126 nd 12 58 11 99 82 67 125 30 31 66 sm 12 12 2 28 23 15 26 9 7 14 eu 12 2 0.9 6 12 7 12 4 3 6 gd 22 14 24 10 7 15 dy 26 22 15 26 10 7 15 ho 4 3 5 2 1 3 er 11 8 14 6 4 8 tm 2 1 2 1 0.6 1 yb 12 5 0.8 14 11 8 13 5 4 7 lu 12 0.6 0.1 2 2 1 2 1 0.6 1 hf 12 12 2 29 9 2 7 2 0.5 4 ta 12 1.4 0.5 4 3 1 3 1 0.4 1 pb 9 11 3 16 3 19 16 3 7 16 1 5 11 th 9 12 2 12 9 2 9 3 9 1 2 13 29 1 6 13 u 12 3 0.7 2.7 0.6 3 3 0.2 4 8 average (ave.), standard deviation (sd), and number of analyses (n). fe2o3* is total iron calculated as fe2o3. granophyre from mcbirney (1989a), larsen & brooks (1994), salmonsen & tegner (2013) and r.b. larsen (2016 personal cummunication). transgressive granophyres from hirschmann (1992). basistoppen granophyre from naslund (1989). melanogranophyre from mcbirney (1989a) and r.b. larsen (2016 personal communication). analysed gabbro from table 13. modelled major element compositions (inverted) for siand fe-enrichment trends (model si and model fe, respectively) from tables 21 and 22, respectively. inverted trace element compositions as modelled in this study (figures 60 and 62). missing trace element data were not given in the original sources. standard deviation are calculated for n > 2. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 68 of 139 geusbulletin.org equilibrium with fayalite and quartz, there are no strong indications that the granophyre interacted strongly with the primocrysts (fig. 14) or that fayalite and hedenbergite formed during final crystallisation (fig. 13). there are, however, some indications that the interstitial granophyre melt may have interacted with the andesine plagioclase in the host to form zoning towards pure albite (fig. 13) and thus may not have been in equilibrium with oligioclase. a further complication is that the proportion of granophyre increases upwards in the stratigraphy towards the sh, which is difficult to reconcile with the notion that granophyre is the residual melt resulting from extreme fractional crystallisation. the interpretation advocated here, based on petrographic and compositional evidence, is that the final differentiated product of the skaergaard intrusion can be represented by two melt compositions: ferrodiorite and granophyre. the ferrodiorites above the uza–uzb boundary (table 18), where fractionation appears to have slowed (fig. 30), represent the basaltic differentiate. these ferrodiorites are made of mixtures of granophyric melt and mafic primocrysts (fig. 14). the former increases in volume upward in the uzc (fig. 15) due to melt migration and accumulation of an immiscible silicic melt. the granophyre component thus represents the final silicic differentiation product, but not necessarily the final differentiation product of the host ferrodiorites. there are no strong indications as to whether ironrich and sodium-rich primocrysts of the ferrodiorites are in equilibrium with the granophyric component. nor are there any clearly observed primocrysts of quartz, orthoclase and albite – the main components of the interstitial granophyre – in the ferrodiorites, although some reactive interactions between the two components cannot be excluded (see section 5.12 for further discussion). 5.5 liquid summation and the liquid line of descent weighted mass balance, liquid summations based on the average compositions of the gabbros zones and subzones, have been used to infer the bulk (or parental) magma composition as well as its liquid line of decent. the calculations require knowledge of the corresponding zone proportions of the intrusion, including the ls, ubs and mbs (nielsen 2004) and average compositions of the respective zone divisions (mcbirney 1989a; hoover 1989b; nielsen 2004; salmonsen & tegner 2013). although the approaches are similar, the details in the calculations may vary between (1) starting with an initial composition and subsequently subtracting the zone compositions (andersen 2006; nielsen et al. 2009) and (2) starting with an end composition, such as a granophyre, and subsequently adding the zone compositions (hunter & sparks 1987; tegner 1997; tegner & cawthorn 2010). furthermore, the modelling often modifies the starting composition by adding a melanogranophyre component (5%; andersen 2006; nielsen et al. 2009) or by adding a melanogranophyre component amounting to 15% in the uzc, 5–8% in uzb, and 2% in the remaining ls (e.g. tegner 1997; tegner & cawthorn 2010). andersen (2006) and nielsen et al. (2009) used an initial bulk-gabbro composition from nielsen (2004) that represented the base of the lza with 5% added melanogranophyre and adjusted to equilibrium with olivine fo68 and the derived melt at the base of mz to be quartz-normative. such adjustments of adding granophyric compositions to the gabbros are equivalent to assuming that the sampled gabbros are not representative of the residual mush and may have lost a component of granophyre (or similar silicic melt), presumably following latestage liquid immiscibility (see section 5.12). fig. 42a compares the previously calculated liquid lines of decent. tegner (1997) and tegner & cawthorn (2010) observed that iron content of plagioclase increases upwards in the ls and relate this to a corresponding increase in the iron content of melt (cf. lundgaard & tegner 2004). their summation calculation assumes an average ferrodioritic rock composition at the centre of uzc (tegner 1997). nielsen et al. (2009) and andersen (2006) use a melanogranophyre terminal composition, while hunter & sparks (1987) assumed a granophyre as the terminal composition (tinden sill). wager (1960) and wager & brown (1967) used a graphical method based on a hz basaltic parental composition (eg 4502; fig. 42a). their results indicated pronounced iron enrichment with modest silica enrichment into the uzb and well past feti oxide appearance (lzc; not shown on fig. 42a) and resulted in a melanocratic granophyre as the terminal differentiation product. consequently, their results for the uz are similar to those of andersen (2006) and nielsen et al. (2009). fig. 42b shows new summation calculations using the gabbro compositions from table 13 and three different silicic end compositions from table 18 (sh ferrodiorite, melanogranophyre and granophyre) shown as model a, b, and c, respectively, without making assumptions about the fate of a granophyric component. other variables in the calculations include zone proportions after nielsen (2004) and the bulk-gabbro compositions of the mbs, ls and ubs from hoover (1989b), mcbirney (1989a) and salmonsen & tegner (2013). the initial magma estimated for each of these is similar (table 19). it is noticeable that all the modelled liquid lines of descent show strong iron enrichment without much silica enrichment throughout the ls and into the uzb (fig.  42b). the appearance of feti oxides (lzc) results in a short reversal in iron enrichment that resumes through the https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 69 of 139 geusbulletin.org a b c d a b c lza, hz lzb lzc mz uza uzb uzc x1 x2 granophyre lzc melanogranophyre sh ferrodiorite granophyre added x1–2 ~lzc nielsen et al. (2009) & andersen (2006) tegner & cawthorn (2010) granophyre melanogranophyre uzc ferrodiorite hunter & sparks (1987) mz eg 4502 sk jj kt39 sc1 cm c j j c 35 45 55 65 75 35 45 55 65 75 2 6 10 14 18 22 26 30 35 45 55 65 75 melt sio2 (wt%) m el t f eo * (w t% ) melt sio2 (wt%) melt sio2 (wt%) a b c a mcbirney & naslund (1990) b this study (thy et al. 2006) c this study (toplis & carroll 1996) d ariskin (2002) experimental modeling: a ferrodiorite b melanogranophyre c granophyre end-member melt: ls sk gabbro mz hanghøj et al. (1995) cm thy et al. (2006) sc1 toplis & carroll (1995) kt39 hoover (1989) sk jj jakobsen et al. (2009) eg 4502 wager & brown (1967) parental melt: j jakobsen et al. (2011) c charlier & grove (2012) immiscible conjugate melt: fig. 42 calculations of the liquid line of descent and initial melt composition shown on feo* vs. sio2 binary diagrams (wt%). a: classic studies by hunter & sparks (1987), andersen (2006), nielsen et al. (2009) and tegner & cawthorn (2010). although the calculations assume or result in three different endmelt compositions (granophyre, melanogranophyre and uzc ferrodiorites), they also assume or result in approximately similar parental melt compositions for lza and hz (see table 19). the calculations by wager (1960) and wager & brown (1967) assume an initial melt composition with much lower initial feo* (eg 4502), but still suggest a liquid line of descent (not shown on diagram) similar to that of nielsen et al. (2009) and andersen (2006). various parental melt compositions, as also suggested, are shown together with the immiscible conjugate melts found by jakobsen et al. (2005, 2011) and charlier & grove (2012). b: summation calculations based on the present ls gabbro compositions (table 13) and end-melt compositions (table 18). the three main liquid lines of descent using different end-melt modelling are: ferrodiorite, melanogranophyre and granophyre. the bulk-gabbro zone compositions are shown for comparison for both ls and the bulk intrusion (sk). the effects of adding granophyre are illustrated (orange curves): the curve marked ×1 uses the equivalent granophyre proportions of tegner & cawthorn (2010) and the curve marked ×2 uses double these proportions. c: experimental and forward experimental modelling. data shown include the melting results of mcbirney & naslund (1990), the two end-member liquid line of descent based on toplis & carroll 1996, thy et al. (2006), tables s1 and s2 and the modelling of ariskin (2002). mz and uza, accompanied by a modest increase in both silica and iron with the latter reaching 23–24 wt% feo*. evolution beyond uza depends on the choice of the granophyric terminal melt (model ‘c’ in fig. 42b) displaying the most marked decrease in iron and increase in silica (fig. 42b). the effects of restoring the gabbro compositions to assumed mush compositions (see section 5.12 for further details), in this case by adding granophyric melt to the gabbro, is also illustrated in fig. 42b by the lines labelled ×1 and ×2. in the former, it is assumed that granophyre was added to uza, uzb, and uzb in the proportions 2%, table 19 summary of calculated parental melt compositions for the skaergaard intrusion using the summation method end-melt gabbro granophyre t & c nielsen andersen (modified a) base of: hz lza hz lza sio2 47.86 47.58 47.90 47.63 tio2 3.17 3.40 3.09 2.99 al2o3 14.52 13.55 13.80 13.38 feo* 14.30 15.78 15.43 15.65 mno 0.21 0.23 0.24 0.23 mgo 6.57 6.16 6.13 7.10 cao 10.18 9.95 10.18 9.87 na2o 2.59 2.63 2.57 2.49 k2o 0.35 0.38 0.40 0.39 p2o5 0.26 0.34 0.28 0.27 mg# 0.45 0.41 0.41 0.45 a composition of this study is modified by adding granophyre as 2% for hz–uza, 5% for uzb, and 15% for uzc. mg#: calculated as mg/(mg+fe) with iron as total fe. t&c: tegner & cawthorn (2010), their table 1. nielsen: nielsen (2004), sk-tfdn in their table 3. andersen: andersen (2006), their table 3. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 70 of 139 geusbulletin.org 5% and 15%, respectively. these amounts are similar to the proportions assumed in the model of tegner & cawthorn (2010), although the composition of granophyre is different. in the second case (×2), the amounts of granophyre are doubled in each subzone. these trapped mush models show marked differences in the liquid line of descent compared to model c without the added granophyre. this exercise illustrates the sensitivity of the liquid line of descent to assumptions about the amount and distribution of trapped granophyre. we submit that this, together with the composition of residual liquid(s), can account for at least some of the diversity of proposed liquid trends from strong iron enrichment at nearly constant silica-to-iron depletion and strong silica enrichment. nevertheless, ascribing a unique liquid line of descent for the skaergaard intrusion based on these considerations alone remains tenuous and highlights a shortcoming of summation models. 5.6 forward modelling of major oxides the systematic cryptic variation in mineral compositions from the start of crystallisation to the final stages has been interpreted to reflect magma differentiation in a closed, or essentially closed, chamber (wager & deer 1939; wager & brown 1967). the essential mechanisms controlling the differentiation are widely accepted as one of solid-liquid fractional crystallisation resulting from an accumulation of mineral components on the base and walls of the chamber and the return of melt to a mainly convective and homogenised magma chamber (maaløe 1976b; mcbirney 1995, 1996; tegner et al. 2009). attempts to model perfect fractional crystallisation for the skaergaard intrusion have been made using forward modelling approaches constrained by relevant experimental phase equilibria (e.g. toplis & carroll 1995, 1996; thy et al. 2006, 2008, 2009a). the success of such forward modelling is measured by its ability to predict essential features of the observed mineral assemblage and cryptic mineral variations, and thereby the composition of the coexisting melts. the summation method (see section 5.5) complements the forward modelling by estimating the liquid line of descent and initial melt composition (fig. 42) using information on the composition of the gabbro zones and subzones in addition to assumptions of the initial or end-stage melt compositions (see section 5.4). 5.6.1 crystallisation conditions and parental melt compositions the forward modelling hinges on the identification of parental melt compositions for the magma chamber and the acceptance of crystallisation conditions. the initial fo2 for the skaergaard magma is believed to be close to the fmq oxygen buffer (morse et al. 1980; toplis & carroll 1995; thy et al. 2006; see section 5.2). but the parental magma composition remains uncertain (table 20). early attempts to constrain this focused on chilled margins of the intrusion (wager 1960, 1961; biggar 1974; mcbirney 1975; naslund 1984; hoover 1989a) and later on an associated dyke swam (brooks & nielsen 1978, 1990; nielsen 1978; toplis & carroll 1995, 1996; thy et al. 2006; jakobsen et al. 2010). examinations of melt inclusions in plagioclase have also suggested primary melt composition at the mz level (hanghøj et al. 1995; jakobsen et al. 2010). the study by jakobsen et al. (2010) on melt inclusions in troctolite inclusions in a skaergaard dyke suggested primary melt compositions with significantly higher mg# than those suggested by the associated dyke compositions and summation calculations (fig. 42b). more recent work has tried to identify units of the contemporaneous east greenland plateau flood basalts as likely sources for the initial skaergaard magma (andreasen et al. 2004; nielsen 2004; nielsen et al. 2019a; cho et al. 2022). a summary of initial skaergaard magmas can be found in table 20. the original chilled margin composition table 20 suggested initial melt compositions for the skaergaard intrusion (wt%)   eg4507 kt39 sc1 cm nielsen 04 sk jj sio2 48.52 50.35 48.75 47.72 47.99 48.82 tio2 1.18 2.65 2.90 2.88 3.04 2.24 al2o3 17.38 13.47 14.89 13.76 13.90 14.42 feo* 9.71 13.46 13.09 15.27 15.13 12.13 mno 0.16 0.20 0.25 0.22 0.18 mgo 8.70 6.57 6.49 6.34 6.30 6.06 cao 11.48 10.15 10.89 10.64 10.18 12.57 na2o 2.39 2.43 2.70 2.46 2.57 3.01 k2o 0.25 0.57 0.30 0.42 0.40 0.38 p2o5 0.22 0.15 0.26 0.27 0.20 mg# 0.614 0.465 0.469 0.425 0.426 0.51 normalised anhydrous to 100% and with all iron as feo. mg#: mg/(mg+fetotal). eg4507: wager (1960). kt39: hoover (1989b). sc1: toplis & carroll (1995) prepared without mno and p2o5. cm: composition used in this study based on brooks & nielsen (1978, 1990). nielsen 04: nielsen (2004). sk jj: jakobsen et al. (2010). https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 71 of 139 geusbulletin.org (eg4507) suggested by wager (1960) was used by biggar (1974) for a one-atmosphere study of the effects of fo2 and temperature on phase equilibria. the chilled margin composition ‘kt39’ was used for low-pressure experimental work by hoover (1989a) and snyder et al. (1993) to study the effects of fo2 and temperature. the composition ‘cm’ that represents the chilled margin composition of an associated dyke is very similar to the bulk skaergaard intrusion calculated by nielsen (2004) and obtained in this study using the summation technique (see section 5.5 and table 19). the latter is the initial skaergaard composition used for the forward modelling of toplis & carroll (1995) and thy et al. (2006). 5.6.2 experimental foundation our forward modelling relies on one-atmosphere melting experiments from liquidus to near solidus conditions using a set of skaergaard-related chilled dyke margins (brooks & nielsen 1978, 1990). the experimental methods, results and forward modelling procedures were presented by thy et al. (2006) and further developed by thy et al. (2009a). here, we only summarise the main experimental results and modelling approach relevant to the present discussion. a total of four olivine and hypersthene cipw normative compositions of dyke margins were used, with mg/(mg+fetotal) = 0.51–0.41, cao/al2o3 = 0.83–0.74 and tio2 = 2.05–4.36 wt%. these compositions are typical of ferrobasalts occurring within the coeval east greenland plateau lavas (brooks & nielsen 1978; larsen et al. 1989; pedersen et al. 1997; tegner et al. 1998; andreasen et al. 2004; nielsen et al. 2019a) and the rift zones of iceland (e.g. thy 1983, 1989). they are thus plausible candidates for the initial melt composition for the skaergaard intrusion. equilibrium melting experiments were conducted at the fmq oxygen buffer using a co-co2 gas mixture. the melting temperatures ranged from the liquidus at 1182–1154°c to the near solidus at 1078–1089°c and produced olivine, plagioclase, augite, ilmenite, magnetite and pigeonite coexisting with melts of mg/(mg+fetotal) down to 0.26–0.22 and with maximum tio2 contents of 4.53–5.20 wt%. cumulative modes are illustrated in fig. 43. plagioclase (an73-60) and olivine (fo77-70) are on the liquidus at high temperature with almost all liquid remaining (f ≈ 1.0), followed by augite (mg# 77–70) on the liquidus at f = 0.90 and coexisting ilmenite and magnetite at f = 0.35. pigeonite replaces olivine for some experiments at f = 0.30 after the appearance of feti oxides. the modal variations as a function of f form the basis for our forward modelling of the liquid lines of decent (fig. 43), with the extrapolated low f values. 5.6.3 fractional crystallisation model the modal variation shown in fig. 43 as a function of f is for equilibrium melting experiments. this means that each melting experiment used in its construction is of constant composition and oxygen content, but variable temperature. as such, the experimental results are directly applicable for modelling equilibrium crystallisation, but not for fractional crystallisation, where the crystallised solid is removed from the system and thus the composition changes with progressive crystallisation and lower temperature. these instantaneous fractional crystallisation modes can be obtained as the derivatives to the equilibrium mode variations. since these latter can be approximated by linear equations, fractional crystallisation can, because the slopes are constants, be modelled by constant modes. using this approach, thy et al. (2006) constrained the plagioclase and olivine cotectic by a constant mode of 26  wt% olivine and 74  wt% plagioclase and the plagioclase, olivine and augite cotectic by 9  wt% olivine, 45  wt% plagioclase and 43  wt% augite. for conditions where pigeonite replaces olivine, the modes were calculated as 19  wt% pigeonite, 41  wt% plagioclase and 40  wt% augite. estimates of the feti oxide modes are complicated by the restrictions in the low-temperature experimental data (thy et al. 2009a). two different hypothetical sets of modes were therefore used to approximate the addition of feti oxides to the fractionating assemblage. the first used the experimental results of thy et al. (2006) and toplis & carroll (1995), suggesting constant feti oxide modes resulting in 8  wt% olivine, 37 wt% plagioclase, 36 wt% augite, 13 wt% ilmenite and 6 wt% magnetite or 34 wt% plagioclase, 33 wt% augite, 16 wt% pigeonite, 11 wt% ilmenite and 6 wt% magnetite, plagioclase olivine augite melt +pigeonite feti oxides 0.4 0.6 0.8 1.00.20 melt fraction remaining (f) a cc um ul at iv e fr ac ti on s 0.4 0.6 0.8 1.0 0.2 0 skaergaard dykes melting experiments one atmosphere, fmq fig. 43 summary of the one-atmosphere melting experiments on skaergaard-related dykes of thy et al. (2006) shown as the accumulative mineral fractions vs. melt fraction remaining (f). short red dashed line: approximate transition from olivine to pigeonite. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 72 of 139 geusbulletin.org respectively, dependent on whether olivine or pigeonite is stable. this first set of feti oxide saturated modes produces a strong si enrichment in the liquid lines of descent similar to that suggested by hunter & sparks (1987) and toplis & carroll (1996). the second set of modes is derived from the assumption that neither ilmenite nor magnetite modes are constant with progressive crystallisation, but rather decrease upward in the stratigraphy as tio2 and possibly fe2o3 decreases. this was suggested by thy et al. (2009a) by showing that both modal contents of ilmenite and magnetite decrease in the gabbros to about 5 wt% ilmenite and < 1 wt% magnetite in the uzc (fig. 15). this latter set of feti oxide saturated modes produces a strong fe enrichment in the liquid line of descent with similarities to those suggested by mcbirney & naslund (1990) and tegner (1997). temperature is determined for olivine or pigeonite saturated melts by a linear correlation between the experimental temperatures and the melt content of mgo as: t = 995.35 + 26.365 × mgo (2) where t is temperature in °c, and mgo is the wt% content of mgo in the melt. the compositions of the crystallising minerals are determined by either single or binary oxide-exchange coefficients either as functions of temperature or as constants. the fe2o3 content of augite is calculated as a constant fraction of total iron (0.18) for the whole ls, while pigeonite fe2o3 is assumed to be zero, and for ilmenite and magnetite the fe2o3 contents are based on the results of toplis & carroll (1995) as functions of δfmq. the fo2 of the melt is calculated from the melt fe2o3/feo ratio and temperature is calculated using the formulation of kilinc et al. (1983). the crystallisation order and zone and subzone divisions are modelled using iterative determinations of best fit to the observed an content of plagioclase and fo content of olivine to constrain the appearances of augite, feti oxides and pigeonite. apatite is not included in the calculations. based on the above assumptions and further details provided by thy et al. (2006), the instantaneous composition of the solid fraction removed during fractional crystallisation can be calculated. the instantaneous melt compositions (or fractional crystallisation) can thus be modelled by solving sequentially the mass-balance equation: ln = (1 – i) × ln+1 + i × sn+1 (3) for the new melt ln+1, knowing the previous melt ln and the new solid sn+1 compositions, in steps from n = 1 to n = ∞ and a constant mass fraction of i for each step. the calculations will approximate perfect fractional crystallisation for i → 0 for which reason a practical value of 0.01 is used in the calculations (equivalent to steps of c. 1°c). 5.6.4 results of forward modelling previous modelling has proposed two contrasting liquid lines of descents, both past feti oxide crystallisation, identified by either si enrichment with restricted fe enrichment (hunter & sparks 1987; toplis & carroll 1996) or by prolonged fe enrichment with restricted si enrichment (mcbirney & naslund 1990; tegner 1997). thy et al. (2006, 2009a) modelled these two contrasting liquid lines of descent using the same initial melt composition (cm), oxygen fugacity at the initial fmq, and silicate phase equilibria (fig. 42c). the differences for the two liquid lines of descent were obtained by changing the modes of the feti oxides in the calculations. the si-enrichment trend was modelled by using the experimental oxide modes, while the fe-enrichment trend was obtained using oxide modes similar to those observed in the intrusion (thy et al. 2006, 2009a). the resultant liquid lines of descent are summarised in supplementary tables s1 and s2 using total calculation steps of 180, solid fraction to 0.80–0.84 (f = 0.20–0.16) and t of 1162–1000°c. the exposed stratigraphic column of the skaergaard ls varies from f = 0.76 at the base of lza to f → 0 at the top of the sh, using the volume interpretation of nielsen (2004) and eq. (1), relating f to stratigraphic height (m; tegner et al. 2009). the stratigraphic height (or depth) for f = 1 is according to eq. (1) at c. 365 m below the exposed base of the ls. to allow a direct comparison with observations, the forward modelling of the experimental melt compositions were broken into the same segments used throughout this study in the data tables. this is done by fitting polynomial curves to the forward modelling as functions of liquid fraction remaining and extrapolated to both f = 1 and f = 0. the results of this curve fitting are variably slightly negative for mgo, cao and tio2 from upper uza to uzb for which small positive values have been adopted for the calculations. the results of the curve fitting are summarised in table 21 for the si-enrichment models and in table 22 for the fe-enrichment models. fig. 44a illustrates the two curve-fitted liquid lines of descent in terms of their feo or feo* contents, both as a function of sio2 (feo* is the total iron calculated as feo). as pointed out by thy et al. (2006, 2009a) and discussed here, the difference in the enrichments for the two basic models are mainly due to differences in the amount of feti oxides removed from the magma, giving rise to the contrasting iron versus silica variations and contrasting evolutions of fo2, as measured by the fe2o3/feo melt ratio (see section 5.2). the forward modelling of perfect fractional crystallisation, using phase equilibria of evolved https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 73 of 139 geusbulletin.org feti basaltic dykes associated with the intrusion, reproduces some aspects of the observed cryptic variation for plagioclase, olivine and clinopyroxene (fig. 44b). the variations reproduced in the modelling is overall similar to the observed cryptic co-variation of the main silicate mineral phases – although some differences are apparent when considered in detail. both the initial and the terminal co-variation of plagioclase and olivine and clinopyroxene fit the observed cryptic mineral compositions. the forward models reasonably reproduce variation in the lz, while the mz and uz results deviate from the observed cryptic co-variation displaying more pronounced convex table 21 modelled silica enrichment melt compositions zone stratigraphic height (m) f t (°c) logfo2 sio2 tio2 al2o3 fe2o3 feo feo* mno mgo cao na2o k2o p2o5 total pl ol pyroxene (mg#) high low (wt%) (wt%) (wt%) (wt%) (wt%) (wt%) (wt%) (wt%) (wt%) (wt%) (wt%) (wt%) (wt%) (an) (fo) ubs 2169 2165 0.00 1012 –13.5 74.48 0.10 10.01 0.12 9.59 9.70 0.29 0.10 1.83 4.08 3.29 0.10 103.99 0.019 0.000 0.000 uzc 2165 2165 0.00 1020 –13.5 74.43 0.10 10.01 0.12 9.61 9.72 0.29 0.10 1.85 4.08 3.28 0.10 103.98 0.020 0.000 0.000 uzc 2165 2164 0.00 1025 –13.5 74.43 0.10 10.01 0.12 9.61 9.72 0.29 0.10 1.85 4.08 3.28 0.10 103.98 0.020 0.000 0.000 uzc 2163 2147 0.00 1021 –13.5 74.40 0.10 10.01 0.13 9.62 9.74 0.29 0.10 1.85 4.08 3.28 0.10 103.97 0.021 0.000 0.000 uzc 2144 2141 0.00 1027 –13.5 74.15 0.10 10.02 0.15 9.72 9.86 0.29 0.10 1.93 4.07 3.26 0.10 103.88 0.026 0.000 0.000 uzc 2091 2081 0.01 1037 –13.4 73.39 0.10 10.04 0.23 10.00 10.21 0.29 0.10 2.13 4.05 3.19 0.10 103.62 0.042 0.000 0.000 uzc 2075 2060 0.01 1036 –13.4 73.16 0.10 10.05 0.25 10.09 10.32 0.29 0.10 2.20 4.04 3.16 0.10 103.54 0.047 0.000 0.000 uzc 2060 2046 0.01 1042 –13.4 72.93 0.10 10.05 0.28 10.18 10.42 0.29 0.10 2.26 4.04 3.14 0.10 103.46 0.052 0.000 0.000 uzb 2044 2030 0.02 1031 –13.4 72.68 0.10 10.06 0.30 10.27 10.54 0.29 0.10 2.33 4.03 3.12 0.10 103.38 0.057 0.000 0.000 uzb 2024 1984 0.02 1043 –13.3 72.37 0.10 10.07 0.33 10.38 10.68 0.29 0.10 2.41 4.02 3.09 0.10 103.27 0.064 0.000 0.000 uzb 1962 1921 0.03 1045 –13.2 71.35 0.10 10.10 0.44 10.76 11.15 0.29 0.10 2.69 3.99 2.99 0.10 102.91 0.085 0.000 0.000 uzb 1902 1862 0.04 1047 –13.2 70.31 0.10 10.12 0.54 11.14 11.63 0.29 0.10 2.98 3.96 2.90 0.10 102.55 0.107 0.000 0.000 uzb 1843 1801 0.05 1052 –13.1 69.24 0.10 10.15 0.65 11.54 12.12 0.29 0.10 3.27 3.93 2.79 0.10 102.16 0.129 0.000 0.000 uzb 1782 1748 0.06 1050 –13.0 68.09 0.10 10.18 0.77 11.95 12.65 0.30 0.10 3.58 3.89 2.69 0.10 101.75 0.152 0.000 0.000 uzb 1728 1691 0.07 1055 –12.9 67.05 0.10 10.21 0.88 12.32 13.12 0.30 0.10 3.87 3.85 2.59 0.10 101.37 0.172 0.000 0.000 uzb 1671 1630 0.09 1055 –12.8 65.93 0.10 10.23 1.00 12.72 13.62 0.30 0.10 4.17 3.82 2.48 0.10 100.96 0.194 0.000 0.000 uzb 1620 1600 0.10 1060 –12.7 64.92 0.10 10.25 1.11 13.08 14.08 0.30 0.10 4.45 3.78 2.39 0.30 100.78 0.214 0.000 0.000 uza 1561 1527 0.11 1063 –12.6 63.73 0.10 10.28 1.23 13.49 14.60 0.30 0.10 4.77 3.74 2.28 2.00 102.02 0.237 0.000 0.000 uza 1506 1431 0.13 1067 –12.5 62.62 0.10 10.30 1.35 13.88 15.09 0.30 0.10 5.08 3.70 2.17 1.87 101.46 0.258 0.000 0.013 uza 1403 1343 0.15 1072 –12.2 60.54 0.17 10.34 1.58 14.57 15.99 0.31 0.10 5.64 3.62 1.97 1.63 100.12 0.297 0.039 0.120 uza 1323 1282 0.18 1072 –12.1 58.96 0.39 10.37 1.75 15.09 16.66 0.31 0.26 6.07 3.55 1.82 1.46 100.00 0.326 0.087 0.199 uza 1268 1237 0.19 1073 –11.9 57.90 0.76 10.39 1.87 15.42 17.10 0.31 0.42 6.36 3.50 1.72 1.35 100.00 0.345 0.120 0.250 uza 1218 1188 0.21 1074 –11.8 56.97 1.09 10.40 1.97 15.70 17.48 0.31 0.58 6.62 3.46 1.63 1.26 100.00 0.362 0.149 0.294 uza 1178 1131 0.22 1075 –11.7 56.22 1.35 10.41 2.06 15.92 17.78 0.31 0.71 6.82 3.42 1.56 1.18 100.00 0.375 0.173 0.327 mz 1098 1055 0.25 1077 –11.5 54.83 1.84 10.44 2.22 16.32 18.32 0.31 0.98 7.20 3.35 1.42 1.05 100.00 0.400 0.219 0.389 mz 1053 1032 0.27 1081 –11.4 54.08 2.10 10.45 2.30 16.52 18.60 0.32 1.14 7.41 3.31 1.35 0.98 100.00 0.414 0.245 0.421 mz 1024 1009 0.28 1083 –11.3 53.61 2.27 10.45 2.36 16.64 18.76 0.32 1.25 7.54 3.29 1.31 0.94 100.00 0.423 0.261 0.441 mz 1003 1002 0.29 1085 –11.3 53.28 2.38 10.46 2.39 16.73 18.88 0.32 1.32 7.63 3.27 1.27 0.92 100.00 0.429 0.272 0.455 mz 988 961 0.29 1088 –11.2 53.05 2.46 10.46 2.42 16.78 18.96 0.32 1.38 7.69 3.25 1.25 0.90 100.00 0.433 0.281 0.464 mz 953 930 0.30 1087 –11.2 52.53 2.64 10.47 2.48 16.90 19.14 0.32 1.51 7.84 3.22 1.20 0.85 100.00 0.443 0.300 0.485 mz 921 900 0.32 1099 –11.1 52.07 2.80 10.48 2.54 17.00 19.29 0.32 1.63 7.97 3.19 1.15 0.82 100.00 0.451 0.317 0.503 mz 890 875 0.33 1093 –11.0 51.65 2.95 10.48 2.59 17.09 19.42 0.32 1.75 8.09 3.16 1.11 0.78 99.99 0.460 0.333 0.520 lzc 847 816 0.35 1095 –10.9 51.09 3.15 10.49 2.65 17.20 19.58 0.32 1.92 8.25 3.12 1.05 0.74 99.99 0.471 0.355 0.541 lzc 808 798 0.36 1093 –10.8 50.62 3.31 10.50 2.71 17.27 19.71 0.32 2.08 8.38 3.08 1.01 0.70 99.99 0.480 0.374 0.559 lzc 784 742 0.37 1096 –10.7 50.34 3.40 10.51 2.74 17.31 19.78 0.32 2.18 8.46 3.06 0.98 0.68 99.99 0.486 0.386 0.569 lzc 723 703 0.40 1103 –10.6 49.70 3.62 10.53 2.82 17.38 19.92 0.32 2.43 8.65 3.00 0.91 0.64 99.99 0.501 0.415 0.593 lzb 703 681 0.41 1099 –10.5 49.51 3.68 10.54 2.85 17.39 19.95 0.32 2.51 8.71 2.99 0.89 0.62 99.99 0.506 0.424 0.600 lzb 634 580 0.44 1108 –10.3 48.91 3.88 10.57 2.92 17.40 20.02 0.32 2.81 8.89 2.92 0.82 0.58 99.98 0.521 0.455 0.622 lzb 558 488 0.47 1110 –10.2 48.38 4.05 10.61 2.98 17.34 20.02 0.32 3.13 9.07 2.86 0.76 0.54 99.98 0.538 0.487 0.642 lzb 447 367 0.53 1113 –9.9 47.82 4.19 10.69 3.04 17.13 19.87 0.32 3.61 9.28 2.76 0.69 0.50 99.97 0.563 0.529 0.667 lzb 346 221 0.58 1121 –9.7 47.54 4.23 10.80 3.06 16.82 19.58 0.31 4.05 9.44 2.68 0.64 0.46 99.96 0.585 0.564 0.685 lzb 177 173 0.67 1126 –9.3 47.46 4.10 11.08 3.00 16.09 18.79 0.30 4.75 9.66 2.56 0.59 0.41 99.95 0.622 0.615 0.714 lza 161 137 0.68 1119 –9.3 47.47 4.08 11.12 2.99 16.01 18.70 0.30 4.82 9.68 2.55 0.59 0.40 99.94 0.626 0.619 0.716 lza 125 107 0.70 1126 –9.3 47.51 4.02 11.21 2.97 15.82 18.49 0.30 4.96 9.72 2.53 0.58 0.39 99.94 0.634 0.629 0.722 lza 96 27 0.71 1131 –9.2 47.54 3.96 11.29 2.94 15.66 18.31 0.30 5.07 9.76 2.52 0.58 0.38 99.93 0.641 0.636 0.727 lza 7 0 0.76 1126 –9.1 47.67 3.77 11.57 2.84 15.16 17.72 0.29 5.39 9.90 2.48 0.56 0.34 99.92 0.660 0.657 0.744 f=1 –365 1.00 1163 –9.0 47.22 2.82 13.80 2.09 13.35 15.23 0.24 6.26 11.24 2.47 0.34 0.30 99.92 0.695 0.733 0.780 melt composition calculated from the silica-enrichment forward modeling by thy et al. (2006) and reasonable values for p2o5 in uzc. feo*: total iron as feo. the modelled tio2 and mgo becomes negative at f ~ 0.13 in the upper part of uza and were fixed at small positive values for the rest of the stratigraphy. t is calculated from plagioclase an (mol%) content based on thy et al. (2009b, 2013). f: fraction of melt remaining. values in italic are assumed. see the supplementary file s3 for complete modeling results. an and fo as mole fractions. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 74 of 139 geusbulletin.org patterns (fig. 44b). the reason for this can be attributed to the effects of fe3+ that may not properly have been included in the fe-mg partitioning or due to the experimental conditions used in the experimental foundation for the forward modelling. what is perhaps more significant is that the observed coexisting mineral compositions (fo0.0 and an~30) at the sh are reproduced already at f = 0.14, corresponding to somewhere in the upper part of the uza (fig. 44b). this can probably be attributed to the assumption of perfect fractional crystallisation and that a component of imperfect fractional crystallisation (equilibrium, in situ fractionation or liquid immiscibility) may be table 22 modelled iron enrichment melt compositions zone stratigraphic height (m) f t (°c) logfo2 sio2 tio2 al2o3 fe2o3 feo feo* mno mgo cao na2o k2o p2o5 total pl ol pyroxene (mg#) high low (wt%) (wt%) (wt%) (wt%) (wt%) (wt%) (wt%) (wt%) (wt%) (wt%) (wt%) (wt%) (wt%) (an) (fo) ubs 2169 2165 0.00 1012 –9.22 55.90 0.92 5.02 8.75 21.21 29.08 0.30 0.40 2.00 2.26 3.92 0.10 100.78 0.040 0.000 0.000 uzc 2165 2165 0.00 1020 –9.22 55.89 0.92 5.04 8.73 21.20 29.06 0.30 0.40 2.00 2.27 3.91 0.10 100.77 0.041 0.000 0.000 uzc 2165 2164 0.00 1025 –9.22 55.89 0.92 5.04 8.73 21.20 29.06 0.30 0.40 2.00 2.27 3.91 0.10 100.77 0.041 0.000 0.000 uzc 2163 2147 0.00 1021 –9.22 55.89 0.92 5.05 8.72 21.20 29.05 0.30 0.40 2.00 2.27 3.91 0.10 100.76 0.042 0.000 0.000 uzc 2144 2141 0.00 1027 –9.26 55.84 0.94 5.13 8.64 21.17 28.95 0.30 0.40 2.00 2.29 3.87 0.10 100.68 0.047 0.000 0.000 uzc 2091 2081 0.01 1037 –9.37 55.71 0.98 5.36 8.41 21.09 28.66 0.31 0.40 2.00 2.35 3.76 0.10 100.45 0.061 0.000 0.000 uzc 2075 2060 0.01 1036 –9.40 55.66 1.00 5.43 8.34 21.06 28.57 0.31 0.40 2.00 2.36 3.72 0.10 100.38 0.066 0.000 0.000 uzc 2060 2046 0.01 1042 –9.43 55.62 1.01 5.50 8.27 21.04 28.48 0.31 0.40 2.00 2.38 3.69 0.10 100.31 0.070 0.000 0.000 uzb 2044 2030 0.02 1031 –9.47 55.57 1.03 5.57 8.20 21.01 28.38 0.31 0.40 2.00 2.40 3.65 0.10 100.24 0.075 0.000 0.000 uzb 2024 1984 0.02 1043 –9.51 55.52 1.05 5.66 8.11 20.97 28.27 0.31 0.40 2.00 2.42 3.61 0.10 100.14 0.081 0.000 0.000 uzb 1962 1921 0.03 1045 –9.64 55.32 1.12 5.96 7.81 20.86 27.89 0.31 0.40 2.00 2.49 3.46 0.10 99.83 0.100 0.000 0.000 uzb 1902 1862 0.04 1047 –9.77 55.12 1.19 6.26 7.51 20.75 27.50 0.31 0.40 2.00 2.56 3.31 0.10 99.50 0.120 0.000 0.000 uzb 1843 1801 0.05 1052 –9.90 54.90 1.27 6.56 7.20 20.64 27.12 0.31 0.40 2.00 2.63 3.15 0.10 99.16 0.140 0.000 0.000 uzb 1782 1748 0.06 1050 –10.03 54.66 1.35 6.88 6.89 20.52 26.72 0.31 0.40 2.00 2.70 2.99 0.10 98.80 0.161 0.000 0.000 uzb 1728 1691 0.07 1055 –10.14 54.43 1.43 7.15 6.61 20.41 26.36 0.31 0.40 2.00 2.76 2.85 0.10 98.46 0.180 0.000 0.000 uzb 1671 1630 0.09 1055 –10.25 54.17 1.52 7.45 6.32 20.29 25.98 0.31 0.40 2.00 2.82 2.70 0.10 98.09 0.201 0.000 0.000 uzb 1620 1600 0.10 1060 –10.34 53.94 1.60 7.70 6.07 20.19 25.65 0.31 0.40 2.00 2.87 2.57 0.10 97.74 0.219 0.000 0.000 uza 1561 1527 0.11 1063 –10.44 53.65 1.70 7.99 5.78 20.06 25.26 0.31 0.40 2.57 2.93 2.41 0.10 97.91 0.240 0.000 0.000 uza 1506 1431 0.13 1067 –10.52 53.36 1.79 8.25 5.52 19.95 24.91 0.32 0.40 3.19 2.98 2.27 1.89 99.93 0.260 0.000 0.000 uza 1403 1343 0.15 1072 –10.65 52.81 1.99 8.71 5.06 19.73 24.29 0.32 0.40 4.29 3.05 2.02 1.65 100.03 0.297 0.014 0.064 uza 1323 1282 0.18 1072 –10.72 52.35 2.14 9.04 4.74 19.56 23.83 0.32 0.42 5.07 3.09 1.83 1.47 100.05 0.324 0.058 0.143 uza 1268 1237 0.19 1073 –10.75 52.03 2.26 9.24 4.54 19.45 23.53 0.32 0.57 5.57 3.11 1.71 1.36 100.16 0.343 0.089 0.196 uza 1218 1188 0.21 1074 –10.76 51.74 2.36 9.42 4.37 19.34 23.27 0.32 0.72 5.98 3.13 1.61 1.26 100.25 0.359 0.117 0.241 uza 1178 1131 0.22 1075 –10.76 51.50 2.44 9.55 4.24 19.26 23.07 0.32 0.86 6.30 3.13 1.53 1.19 100.31 0.372 0.140 0.277 mz 1098 1055 0.25 1077 –10.75 51.02 2.61 9.77 4.02 19.09 22.70 0.32 1.13 6.85 3.13 1.38 1.05 100.40 0.396 0.186 0.343 mz 1053 1032 0.27 1081 –10.73 50.75 2.71 9.89 3.91 18.99 22.50 0.32 1.30 7.13 3.13 1.31 0.98 100.43 0.410 0.211 0.377 mz 1024 1009 0.28 1083 –10.71 50.58 2.77 9.96 3.84 18.92 22.38 0.32 1.42 7.30 3.12 1.26 0.94 100.45 0.418 0.228 0.399 mz 1003 1002 0.29 1085 –10.70 50.46 2.82 10.00 3.80 18.88 22.29 0.32 1.50 7.41 3.12 1.23 0.91 100.45 0.424 0.240 0.414 mz 988 961 0.29 1088 –10.68 50.37 2.85 10.03 3.77 18.84 22.23 0.32 1.56 7.49 3.11 1.21 0.89 100.46 0.428 0.249 0.424 mz 953 930 0.30 1087 –10.65 50.16 2.93 10.10 3.70 18.76 22.09 0.32 1.71 7.66 3.10 1.16 0.85 100.46 0.438 0.269 0.448 mz 921 900 0.32 1099 –10.62 49.98 3.00 10.16 3.65 18.69 21.97 0.32 1.85 7.80 3.09 1.12 0.81 100.47 0.447 0.287 0.468 mz 890 875 0.33 1093 –10.58 49.80 3.06 10.21 3.60 18.62 21.85 0.32 1.98 7.93 3.07 1.08 0.78 100.46 0.455 0.304 0.487 lzc 847 816 0.35 1095 –10.52 49.56 3.16 10.27 3.54 18.51 21.70 0.32 2.18 8.09 3.05 1.03 0.74 100.45 0.467 0.328 0.511 lzc 808 798 0.36 1093 –10.47 49.35 3.24 10.32 3.49 18.41 21.56 0.32 2.36 8.23 3.03 0.99 0.70 100.44 0.477 0.350 0.532 lzc 784 742 0.37 1096 –10.43 49.23 3.29 10.35 3.47 18.35 21.47 0.32 2.47 8.30 3.01 0.96 0.68 100.43 0.483 0.363 0.544 lzc 723 703 0.40 1103 –10.33 48.92 3.41 10.42 3.41 18.19 21.26 0.32 2.77 8.46 2.97 0.91 0.63 100.41 0.498 0.395 0.573 lzb 703 681 0.41 1099 –10.29 48.82 3.45 10.44 3.40 18.13 21.19 0.32 2.86 8.51 2.95 0.89 0.62 100.40 0.503 0.405 0.581 lzb 634 580 0.44 1108 –10.16 48.51 3.58 10.49 3.35 17.93 20.95 0.32 3.21 8.66 2.90 0.84 0.58 100.37 0.519 0.440 0.608 lzb 558 488 0.47 1110 –10.01 48.20 3.72 10.55 3.31 17.69 20.67 0.32 3.59 8.81 2.83 0.79 0.54 100.34 0.537 0.476 0.634 lzb 447 367 0.53 1113 –9.79 47.83 3.88 10.64 3.26 17.30 20.23 0.32 4.15 9.00 2.74 0.73 0.50 100.34 0.562 0.525 0.664 lzb 346 221 0.58 1121 –9.60 47.59 3.99 10.75 3.20 16.90 19.78 0.31 4.65 9.18 2.66 0.68 0.46 100.38 0.585 0.564 0.686 lzb 177 173 0.67 1126 –9.32 47.41 4.06 11.05 3.06 16.12 18.87 0.30 5.44 9.58 2.54 0.61 0.41 100.58 0.623 0.617 0.717 lza 161 137 0.68 1119 –9.30 47.41 4.06 11.09 3.04 16.04 18.77 0.30 5.51 9.63 2.53 0.60 0.40 100.60 0.627 0.622 0.719 lza 125 107 0.70 1126 –9.25 47.41 4.05 11.18 2.99 15.85 18.54 0.30 5.67 9.73 2.52 0.58 0.39 100.67 0.635 0.631 0.725 lza 96 27 0.71 1131 –9.21 47.41 4.03 11.26 2.95 15.70 18.35 0.30 5.79 9.82 2.50 0.57 0.38 100.72 0.641 0.638 0.730 lza 7 0.76 1126 –9.13 47.47 3.93 11.57 2.81 15.21 17.74 0.29 6.14 10.12 2.48 0.53 0.34 100.89 0.660 0.658 0.743 f = 1 –365 1.00 1163 –8.90 47.67 2.53 13.70 2.22 13.22 15.22 0.25 6.36 10.53 2.42 0.44 0.28 99.62 0.696 0.739 0.787 melt composition calculated from the silica-enrichment forward modelling by thy et al. (2006) and reasonable values for p2o5 in uzc. feo*: total iron as feo. the modelled tio2 and mgo becomes negative at f ~ 0.13 in the upper part of uza and were fixed at small positive values for the rest of the stratigraphy. t is calculated from plagioclase an (mol%) content based on thy et al. (2009b, 2013). f: fraction of melt remaining. values in italic are assumed. see the supplementary file s3 for complete modelling results. an and fo as mole fractions. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 75 of 139 geusbulletin.org able to significantly delay fractionation to extreme values at f ≈ 0. a somewhat similar effect was seen in the cryptic variation (fig. 30) that suggests fractionation ceases in the upper part of uzb. finally, there are few, if any, observable differences in the cryptic mineral variation of the two modelled liquid lines of decent, except at low values of f (fig. 44b). this means that studies of the silicate cryptic mineral variation and the phase layering are unlikely to offer definitive conclusions as to the nature of the liquid lines of descent for the intrusion. the diverse proposed melt evolutions for the skaergaard intrusion abundantly attest to this observation (wager & brown 1967; hunter & sparks 1987; mcbirney & naslund 1990; toplis & carroll 1996; nielsen 2004; andersen 2006; thy et al. 2006; tegner & cawthorn 2010). 5.7 additional phase-equilibria constraints basaltic melts, such as the parental melt to the skaergaard intrusion, followed crystallisation and fractionation paths controlled by phase equilibria between melt and silicate and oxide minerals. however, the sampled gabbros represent the accumulated mixtures of variable proportions of solid minerals and a component representing the final trapped melts that only record the evolution of the parental melt and the phase equilibria indirectly. early work used phase equilibria in simplified or haplobasaltic systems to understand the evolution of the skaergaard and other layered intrusions (osborn 1959, 1979; presnall 1966; biggar 1974; maaløe 1976b). later studies expanded the approach to multicomponent basalt systems (irvine 1970a, 1979; jackson 1970; ford 1981). we here use the same approach by displaying the skaergaard gabbro compositions graphically as projections of the pseudo-quaternary basalt system (fig.  45) using the plagioclase, olivine, quartz and diopside (di) components – the latter as a proxy for clinopyroxene (augite). this approach allows important principal observations to be made about crystallisation orders and the dominate mode of differentiation. 5.7.1 projection schemes the analysed gabbros (table 13) and silicate minerals (tables 3–5, 8) are recast using a cipw normative compositions (molecular norm) for plotting. the fe2o3/ feo ratios for the bulk rocks were determined by wet chemistry, while all silicate minerals are assumed to be fe2o3-free. the latter is a simplification for clinopyroxene, but we opted to do this given the difficulties of accurately constraining fe+3 content from stoichiometry based on microprobe data (fig. 22; see section 4.3.3). increasing the amount of ferric ion in clinopyroxene will shift these slightly towards higher normative quartz (q), sio2 (wt%) fe o * or f eo forward modelled melt compositions f = 0 f = 1 f = 0 f = 0 feo feo feo* feo* fe enrichment si enrichment uzb uzb 10 15 20 25 30 45 50 55 60 70 7565 a observed and modelled cryptic variation experimental ol-pl b f = 1 f = 0.14 plagioclase (an mol%) 80 70 60 50 40 30 20 o liv in e (f o m ol % ) o r c lin op yr ox en e (m g/ (m g + fe ) c at io n % ) 80 70 60 50 40 30 20 10 0 clinopyroxene olivine fe enrichment si enrichment modelled cryptic variation f = 1: fo73.9 an69.6 fig. 44 forward modelling of the liquid line of descent and the cryptic mineral variation for the preferred, initial skaergaard magma based on thy et al. (2006, 2009b). a: melt composition of feo or feo* as a function of sio2 (wt%). the two end-member models (fe or si enrichment) are based of the solutions of thy et al. (2006) and toplis & carroll (1996), respectively, extrapolated to f = 1 and f = 0 (f: melt fraction remaining). feo* is total iron calculated as feo. the difference between the curves for feo and feo* reflects the amount of fe2o3. b: olivine (fo mol%) or clinopyroxene (mg/(mg + fe) cation %) compositions as a function of plagioclase (an mol%). data shown are from tables 3–5 with additional primitive gabbro compositions from thy et al. (2008, table 1). iron in clinopyroxene is calculated as total iron. solid curves: forward modelling, compared to the observed coexisting minerals in the ls. the experimental ol-pl is the coexisting olivine and plagioclase at the liquidus in the one-atmosphere melting experiment used to formulate the forward model (thy et al. 2006). green box: extrapolated coexisting plagioclase and olivine at f = 1. the forward modelling fails to reproduce the observations in uzc due to early modelled depletion of mgo to zero and is thus not shown past f = 0.16. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 76 of 139 geusbulletin.org without changing the principal conclusions based on the fe2o3-free assumption. our projection scheme is similar to that used by presnall et al. (1979) where forsterite + fayalite = olivine (ol), anorthite + albite = plagioclase (pl), and diopside (di) + hedenbergite = di. enstatite and ferrosilite are combined as orthopyroxene (opx) and plot midway between ol and q. the projections are from minor components like magnetite, ilmenite, and apatite. this means that to the extent these minor components conform to the simplified cipw normative compositions, they will have no effect on the variability in the quaternary diagram. three triangular projections of the quaternary are used to facilitate a three-dimensional understanding: the di-ol-q (projected from pl), the pl-ol-q (projected from di), and the di-pl-ol (projected from q). the compositions used in this discussion are, as shown in the tables, not individual analyses, but averages of three consecutive samples from the stratigraphy. 5.7.2 crystallisation order and zone divisions the normative compositions of the lz gabbros plot in the quaternary close to the di-pl-ol join, being olivine and hypersthene normative, without reaching into the quartz-normative volume (fig. 45). the uz gabbros (not shown) plot in the same general area as lzc–mz, but with only uzc extending into the q-normative volume. the pl/ol ratios (1.39 ± 0.28) are restricted compared to the more variable di/pl ratio that systematically increases upward from 0.32 ± 0.03 in lza to 0.59 ± 0.12 in lzb and 0.93 ± 0.32 in lzc (fig. 45). in the ternary projections, lzc, mz, and uza gabbros cluster around the ol-normative extension of the modelled liquid lines of descent (fig. 46a) as would be expected for solid crystallisation products from a ferrobasaltic melts similar to those used for the present forward modelling (thy et al. 2006, 2009a). they also cluster around the olivine and hypersthene normative segment of the experimental plagioclase, clinopyroxene and olivine multi-saturated cotectic (marked l(pl,cpx,ol) in fig. 46a) as expected for gabbros formed from crystallisation of these three silicate primocrysts from a large range of melt compositions (thy et al. 2008). the lza gabbros with a low normative diopside content (figs 46a, 47a) are consistent with early accumulation of primocrysts of plagioclase and olivine. this relationship is consistent with the zonal division based on petrography, where wager & brown (1967) and mcbirney (1989a) defined the base of the lzb by a reduction in the amount of oikocrystic augite and appearance of granular augite (see fig. 5). the transition occurs at a height of about 174 m (normalised to the stratigraphic scale adopted in this study, fig. 1) and closely conforms to the marked increase in normative diopside. the lzb and lzc gabbros reflect an upward increasing accumulation of clinopyroxene with lzc eventually reaching a proportion of clinopyroxene predicted from experimentally-based modelling (fig. 47a; thy et al. 2006). this order of lz crystallisation supports the observed zonal division, suggesting early concurrent crystallisation of plagioclase and olivine in lza, followed by clinopyroxene in lzb, and finally adding feti oxide minerals (magnetite and ilmenite) in lzc, where the effects of the feti oxides are not, or only marginally, detected in the projection (fig. 47a). because the part of the ls sampled in this study does not reach into the lowermost hz, the question of whether plagioclase or olivine were the initial crystallising silicate primocrysts cannot directly be resolved. evidence from cambridge drill core i suggests that the hz gabbros, extending for at least 150 m and possibly more, below the exposed part of lza, are similar to the upper part of the lza (maaløe 1976a; holness et al. 2015). holness et al. (2015) suggested that the skaergaard magma was filled by several early pulses of plagioclase and olivine-bearing magma eventually leading to a ballooning of the chamber to its final size. this is consistent with the observation by ariskin (2003) that the initial magma contained plagioclase and olivine phenocrysts. the arrival of a new primocryst, like augite, is expected to result in a flattening of the liquidus and consequently in the increase in proportions of crystallisation (wyllie 1963). an additional effect is an intermittent increase in di ol pl q lza lzb lzc–mz opx lz–mz lzc–mz lzb lza fig. 45 perspective view of the lz–mz gabbros in the quaternary olivine (ol) – quartz (q) – plagioclase (pl) – diopside (di) with projections onto the di–pl–ol and pl–ol–di triangles. a molecular equivalent norm, calculated based on the cipw norm procedure, was used for the projection with ferric and ferrous iron as analysed. the gabbro compositions from table 13 are grouped as lza, lzb and lzc–mz. for simplicity, the uz is not shown. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 77 of 139 geusbulletin.org latent heat release (holness et al. 2007a, b, 2009) that is manifested in an increase in textural grain-boundary maturity effective until the excess heat dissipates. holness et al. (2015) examined the lower part of the intrusion in detail, including the hz, without being able to pinpoint equivocally the lza–lzb boundary based on textural criteria like dihedral angles and maturity of the growth of augite in interstitial melt – although they did proposed a progressive transition between 100 and about 200 m depth in the cambridge drill core. thus, it remains to be explained why the lzb gabbros occupy an intermediate position in the projections between lza and lzc, when lzb and lzc could be expected to have similar olivine-plagioclase-clinopyroxene di q pl ol opx uzc uza di q (pl) opx uzb pi uz oliv +pi aug l (pl,cpx,ol) mol% ol uzc uza uzb ba lza aug mz– lzc lzb di qol (pl) opx lz–mz +pi oliv pi l (pl,cpx,ol) fe-enrichment si-enrichment mol% uza aug hr, wü hr, wü uza uza lzc–mz lza lzb fig. 47 projections from plagioclase (pl) of the quaternary ol–q–pl–di on the triangular di–ol–q diagram (see quaternary diagram, inset) in mol%. data, projection methods and abbreviations in fig. 45. the gabbro compositions are grouped as lza, lzb, lzc–mz, uza, uzb and uzc. a: projection for lz– mz. b: projection for uz. l(pl,cpx,ol) represents melt compositions coexisting with plagioclase, clinopyroxene, and olivine at variable temperatures (thy & lofgren 1992, 1994). +pi indicates the appearance of pigeonite (thy et al. 1998, 1999). the two curves marked ‘fe enrichment’ and ‘si enrichment’ are the liquid lines of descent, which both terminate at the end of uza (thy et al. 2006, 2009a; tables 21 and 22). approx. location of the join ol–pi–aug is marked by the grey shaded triangle. the general effect of non-quaternary normative components in magnetite (hercynite, hr, and wüstite, wü) is indicated by arrows directed away from q. di ol (q) lza lza aug lzb lzc– mz uzc uza di pl ol (q) uzb uzc–b lza– uza lzb lzc–mz uzb uza uzc lz–mz uz di q pl ol opx oliv oliv orthopyroxene pigeonite ferrobustamite aug lza– uza a b (mol%) (mol%)pl plag plag l(pl,cpx,ol) orthopyroxene /pigeonite fig. 46 projections from quartz (q) of the quaternary ol–q–pl–di on to the triangular di–pl–ol diagram (see quaternary diagram, insert) in mol%. a: projection for lz–mz. b: projection for uz. projection scheme and abbreviations in fig. 45. the gabbro compositions from table 13 are grouped as lza, lzb, lzc–mz, uza, uzb and uzc. gabbro is calculated with ferric and ferrous iron as analysed, while all minerals (aug: augite, plag: plagioclase, oliv: olivine, orthopyroxene/pigeonite and ferrobustamite) are calculated with all iron as ferrous iron. average mineral compositions are from tables 3–5, 8 for: plagioclase and clinopyroxene (aug; grouped as lza–uza and uzb–uzc), ferrobustamite, orthopyroxene, pigeonite and olivine. approximate location of the join of pl–ol–aug is marked by the grey shaded triangle. the black square labelled l(pl,cpx,ol) is the piercing, one-atmosphere, plagioclase-augite-olivine cotectic based on melting experiments of thy & lofgren (1992, 1994) and thy et al. (1998, 1999). https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 78 of 139 geusbulletin.org proportions (figs  46a, 47a). there are several explanations although none of these are conclusive. the first explanation is that the lzb parental melts were not saturated in augite, despite the petrographic indications, but, like lza, were only saturated in olivine and plagioclase. the augite component in this case may have originated from a crystallised component of trapped melt estimated from 40 to 10% upward (tegner et al. 2009) in an olivine and plagioclase mush and not in a mush of augite, olivine and plagioclase. modelling the solidification of such a mush for the high melt content, assuming a fractionation path similar to that for the entire ls and near augite saturation in the trapped melt (thy et al. 2006), results in a felsic gabbro at the base of lzb composed of 20 wt% olivine, 65 wt% plagioclase and 15 wt% augite (on a weight basis) plus lesser amounts of orthopyroxene and feti oxides. a calculation assuming closed equilibrium crystallisation (thy et al. 2006) predicts rather similar modes of 15% olivine, 68% plagioclase and 12% augite. thus, despite demonstrating that substantial augite can be formed from the trapped melt, both models fall well short of predicting the 28% augite found in the lzb (table 1). furthermore, the amount of trapped melt decreases substantially upward to a quarter of that from the base of lzb (tegner et al. 2009) and consequently, the modelled modal augite reduces proportionately, assuming cotectic plagioclase and olivine saturation in the main magma. a second, or alternative, explanation is that the trappedmelt content in the lz may be underestimated, thus allowing for a higher effect of interstitial crystallisation of clinopyroxene. however, required amounts by far exceed possible estimates in this study (see section 5.12.6). the texture of the lzb suggests co-crystallisation of plagioclase, olivine, and augite and that the gabbros were controlled by liquids, in turn controlled by these three silicate phases. this implies a cotectic shift towards lower augite modes, relative to the univariant cotectic shift determined for low-pressure north atlantic lavas – including east greenland lavas and skaergaard-related dykes (fig. 47a; thy & lofgren 1992, 1994; thy et al. 1998, 1999, 2006, 2008). such a shift was not detected for skaergaard-related dykes and therefore not included in the experimental data used to forward model the skaergaard magma. cotectic shifts for low-pressure conditions have been suggested for olivine-normative compositions. in particular, works by walker et al. (1979) and presnall et al. (1979) have suggested a marked increase in normative plagioclase in olivine-hypersthene normative melts approaching the ol-pl-di join. later experimental work covering the full melt range of nepheline to hypersthene normative compositions has, however, been unable to support significant cotectic deviations from linearity like that shown in fig. 47 (sack et al. 1987). we suggest that neither differences in the trappedmelt content nor shifts in the olivine-plagioclase-augite cotectic offer plausible explanations for the low modal pyroxene-content in the lzb, and so other explanations must be sought. there is a systematic difference in the texture of lzb, compared to the subzones below and above, namely a general increase in average grain size concurrently with a change in the pyroxene fabric from an oikoclastic fabric in lzb to a dominating granular fabric in lzc (cf. fig. 5). it is possible that this textural variation can be related to variation in crystallisation conditions across the lz and that this could be seen as a suppression of augite crystallisation concurrently with an increase in grain size in the lzb. the variation in crystallisation conditions such as pressure and water activity are expected to affect the relative saturation of augite and other silicate phases. increasing dry pressure is expected to raise augite saturation at the expense of olivine and plagioclase (stolper 1980; grove & kinzler 1992; berndt et al. 2005; whitaker et al. 2007) but this is not a realistic explanation (see section 5.3). the effects of increasing water activity in the melt are known to depress the appearance of augite at the expense of olivine and plagioclase in basaltic melts (almeev et al. 2007, 2012; botcharnikov et al. 2008) and probably also to increase the grain size. however, the skaergaard gabbros lack any compositional and mineralogical evidence to suggest that they originated from anything other than nominally dry melts (see section 5.9). finally, the low augite proportions in lzb could be attributed to a biased sampling procedure, resulting in gabbros that do not represent the original mush composition, perhaps due to the relatively coarser grain size compared to the zones above. however, in view of the sampling results for both lza and lzc and the systematic sampling of the felsic lzb gabbros, such a sampling bias is not considered plausible. 5.7.3 clinopyroxene crystallisation clinopyroxene occupies an elongated volume close to the di-ol join (fig. 46). when projected from pl onto the triangle di-ol-q, as shown in fig. 47 and in more detail in fig. 48, a systematic variation with decreasing diopside content in the lz and mz can be seen as a function of stratigraphy in the ls (fig. 48a). this decrease is followed by an increase upward in the uz (fig. 48b). this variation mirrors that seen in the diopside content of the pyroxenes forming a keel at the boundary between the mz and uza (figs 19, 21). the result is that clinopyroxene plots with decreasing normative diopside up through the stratigraphy until the mz. thereafter the trend reverses to increasing diopside forming a narrow v-shaped pattern (fig. 48b). as pointed out by coombs (1963), this pattern is also reflected in the original https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 79 of 139 geusbulletin.org skaergaard pyroxene analyses on separated minerals by brown et al. (1957) and brown & vincent (1963; fig. 48c). incorporating fe3+ in the calculations will have an insignificant effect. variation in the normative pyroxene composition leads to an overall increase in modal clinopyroxene content upward in the stratigraphy throughout the lz to mz, followed by a corresponding drop in the uz. the result is a dramatic change in the slope of the fractionation path at the mz–uza boundary, accelerating fractionation towards increasing normative q (fig. 49). this is seen on many binary and ternary variation diagrams illustrating the mineral compositions and the melt evolution throughout this study. it is an interesting consequence that while the melt composition is dominantly quartz-normative, the resulting gabbros are consistently olivine and hypersthene-normative. 5.7.4 feti oxide crystallisation the minor normative components apatite, ilmenite, hematite and magnetite were excluded from the projection scheme on the assumption that their compositions are nearly identical to the normative components. this is true for natural apatite and for ilmenite for which hematite is calculated as a separate normative component. however, the norm calculation for natural magnetite assumes ideal magnetite (feo.fe2o3) and in effect calculates ulvöspinel (2feo.tio2) as ilmenite (feo.tio2) and hematite (fe2o3). an occasional small amount of normative rutile (tio2) is ignored. while ilmenite is close to the ideal composition (with hematite being part of the ilmenite solid solution), the magnetite of the skaergaard intrusion (table 9) has an average normative composition of 84 wt% magnetite, ilmenite and hematite with the remaining made up of 16 wt% excess wüstite (feo) and corundum (al2o3). the latter two components can be recast as 1 wt% hercynite (hr, feo.al2o3) and 15 wt% di ol (pl) di q pl ol opx pigeonite a (mol%) augite mz lza lzb lzc lz a lz b m z– lz c di (pl) opx oliv augite uza uzb uzc uzc uzb uz di bustamite q (pl) pigeonite augite uzb,c ubs lz–uza brown et al. (1957) brown & vincent (1963) fe3+/fe2+ as analysed uza b c lz–mz uzc uzb uzamz lza lzb lzc lza mz pigeonite fig. 48 pyroxene variation as a function of ls zones and subzones in the plagioclase (pl) projection on the ol–di–q triangle of the quaternary ol–q– pl–di (see quaternary diagram, inset) in mol%. abbreviations in fig. 45. the average pyroxene compositions from tables 5, 7 and 8 are shown for each zone and subzone of the ls. a: projection for lz–mz. b: projection for uz. c: projection for skaergaard pyroxene analyses on separated minerals by brown et al. (1957) and brown & vincent (1963), assuming either all iron as ferrous iron (dark grey) or as originally analysed (light grey). approximate variation in the locations of the join of olivine–pigeonite–augite is marked for the ls subzones by grey shaded triangles. ubs: upper border series. uzc uzb uza mz lzc lzb lza 0.4 0.5 0.6 0.7 0.8 0.9 0.4 0.3 0.2 0.1 0.0 normative di/(di + ol + q) pyroxene n or m at iv e d i/ (d i + o l + q ) m el t m od el fig. 49 normative fraction of diopside (di) of the modelled liquid line of descent (normative di/(di + ol + q) melt model); si-enrichment trend; table 21) vs. the same fraction for observed clinopyroxenes (normative di/(di + ol + q) pyroxene). average values for the individual zones and subzones of the ls are shown. abbreviations in figs 2, 45. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 80 of 139 geusbulletin.org wüstite (wü, feo). the effect of these non-quaternary components will be to shift the position of the feti oxide-bearing cumulates away from q in the triangular diagram projected from pl (thy 1995; fig. 47a). this non-quaternary shift may affect lzc, the mz and uza, in decreasing magnitude determined by their amounts of magnetite, relative to lza and uzb. this effect must be considered when interpretating the details of the normative projections, but is not large enough to significantly affect the principal interpretation of the present phase equilibria, except for a minor shift of the lzc–mz gabbros away from q compared to the uza gabbros (fig. 47a). 5.7.5 orthopyroxene crystallisation the status of orthopyroxene in the lower parts of the ls has been viewed as either a primocryst (wager & deer 1939; wager & brown 1967; irvine et al. 1998) or a late phase-replacing olivine and clinopyroxene (mcbirney 1989a; mcbirney & naslund 1990; tegner et al. 2009). the latter would include crystallisation of orthopyroxene replacing olivine from the trapped melt (mcbirney 1989a), subsolidus lamellae exsolution and coarsening from high-ca clinopyroxene (lindsley & andersen 1983) or both. the projections from plagioclase onto the triangle ol-di-q (which includes opx) have some bearing on this discussion (fig. 50). the existence of a reaction relationship between olivine primocrysts and melt to form pigeonite has been demonstrated experimentally for relevant skaergaard parental melts (hoover 1989a; snyder et al. 1993; thy et al. 1998, 1999, 2006) to occur at c. f = 0.3, approximately coinciding with the lzc–mz boundary. based on these experiments, the reaction relationships are +pi di q pl ol opx di qol (pl) opx si-enrichment uzb (mol%) lza lzb uza troctolite trapped meltlzc-mz lzb aug di qol (pl) uzb (mol%) lza lzb uza gab bro aug lzc, mz g ab br o lzc lzb di qol (pl) uzb (mol%) lza lzb uza gab bro aug uza lzc-mz mz mz lzc b c d l (pl,cpx,ol) +pi di qol (pl) opx si-enrichment uzb (mol%) lza lzb uza troctolite trapped meltlzc-mz lza aug g ab br o lza a l (pl,cpx,ol) +pi opx l (pl,cpx,ol) +pi l (pl,cpx,ol) lzc-mz opx si-enrichment trapped melt si-enrichment trapped melt fig. 50 summary interpretations of projections from plagioclase (pl) of the quaternary ol–q–pl–di on the triangular di–ol–q diagram (see quaternary diagram, inset) in mol%. a: projection for lza. b: projection for lza. c: projection for lzc and mz. d: projection for uza. data, projection methods and abbreviations in fig. 45. outlines of the zone and subzones of the ls are shown and interpreted in terms of modal gabbro compositions on the join of ol–aug and trapped melt composition on the si enrichment liquid line of descent. l(pl,cpx,ol) represents melt compositions coexisting with plagioclase, clinopyroxene, and olivine at variable temperatures (thy & lofgren 1992, 1994). +pi indicates the appearance of pigeonite (thy et al. 1998, 1999). https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 81 of 139 geusbulletin.org approximated by 0.9 liquid + 0.1 olivine = 0.2 pigeonite + 0.4 plagioclase + 0.4 clinopyroxene (see thy et al. 2006). thus, at the level of the lza, closed crystallisation of the trapped melt estimated by tegner et al. (2009) could result in c. 5–6% pigeonite. this would necessarily have decreased markedly upward as a function of the decrease in the total amount of trapped liquid. the saturation of pigeonite in equilibrium melting experiments occurs at an invariant reacting point in quartz-normative melts (marked ‘+pig’ in fig. 50) after the saturation of feti oxides. however, the corresponding pigeonite saturation point during fractional crystallisation is uncertain, and judging from the present forward modelling, it may occur in less quartz-normative melts in the mz. these considerations point to the importance of pigeonite crystallisation (later inverted to orthopyroxene) during solidification from the interstitial melt of the mush. orthopyroxene exsolution and coarsening of lamellae are also evident in the pyroxenes formed during cooling at near solidus to subsolidus conditions and may have affected the orthopyroxene content. the similarity between the bulk clinopyroxenes analysed by brown et al. (1957) and brown & vincent (1963) and those obtained using the broad-beam technique with an electron microprobe (fig. 19) suggests that exsolution may have had some, but limited effect, on the pyroxene proportions. it thus appears reasonable to reduce the gabbro compositions to the ratio between olivine and clinopyroxene after adding orthopyroxene to olivine as in fig. 50. using this procedure, the average mz–uzc gabbro is calculated as containing 78 mol% clinopyroxene on the ol-aug join (fig. 50c). this compares well to the intersection with the known three-silicate phase co-saturate in basaltic melt (76% clinopyroxene) but is very different if orthopyroxene is not added (85% clinopyroxene). the composition of the lza gabbros can thus be explained by a mixture of olivine (and plagioclase) and trapped melt and does not require primocryst orthopyroxene (fig. 50a). this is consistent with the observation that primocrysts in the lza were olivine and plagioclase. the modal orthopyroxene content of the gabbros (tables 1 and 2) is thus mostly due to interstitial melt crystallisation. the lzb gabbros can likewise be interpreted as a mixture of olivine, plagioclase, augite and interstitial melt (fig. 50b) and does not require orthopyroxene in excess of that crystallised from the interstitial melt – although the low augite content is still unresolved. for the remaining stratigraphic sequence, including the mz, there is likewise no evidence that orthopyroxene exceeds the amount that could have crystallised from the interstitial melt. the mode of orthopyroxene in the mz can be estimated from this study as an average of 6% pigeonite, varying between 3% and 8% (table 2). the modal content of orthopyroxene increases at the expense of olivine upward from the central part of the mz (figs 9, 15). such amounts are similar to the bulk-gabbro content of the lz, but far from the expectations for pigeonite (c. 15%) being a liquidus phase in the mz based on the existing experimental information on skaergaard-like basalt compositions (hoover 1989a; snyder et al. 1993; thy et al. 1999, 2006). the available experimental evidence points to a hypersthene-normative parental melt where olivine is replaced by orthopyroxene during crystal fractionation (grove & bryan 1983; grove & baker 1984; longhi & pan 1988; grove & juster 1989; veksler 2009). this suggests that the parental skaergaard melt was hypersthene (and olivine) normative and fractionated towards an increasing quartz-normative composition (toplis & carroll 1995, 1996; thy et al. 2006). what is less clear from the petrological evidence is, however, whether pigeonite was on the liquidus in the mz. the alternative view, consistent with observations, is that pigeonite was a near-liquidus phase, crystallising from an interstitial melt throughout most of the lower part of the intrusion and never reached the liquidus prior to the interruption of fractional crystallisation by liquid immiscibility in the uzb. 5.7.6 constraints on trapped-melt compositions the gabbros analysed here can be evaluated in the normative projections as mixtures of primocrysts (olivine, plagioclase and pyroxenes) and trapped melt, now represented by a mixture of the same three minerals. a simplified interpretation is shown in fig. 50 as projections from plagioclase onto the triangular ol-pl-q diagram. the evaluations are based on the modelled liquid lines of descent showing si enrichment (table 21). however, using the fe-enrichment trend (table 22) will not significantly change the interpretation (compare with fig. 47a). average gabbro compositions in the subzones are projected to the ol-aug join using the modal compositions of table 2 with orthopyroxene added to olivine and thus assuming that all orthopyroxene is a reaction product. to do this, we treat gabbro compositions as a mixture of olivine, augite and trapped-melt composition obtained from our forward modelling. the three-component system is then finally reduced to a binary system along the ol-aug join and the proportions of primocrysts and trapped melt are calculated using the lever rule (fig. 50). this exercise shows that the composition of lza can be accounted for as a mixture c. 40% olivine gabbro and c. 60% trapped melt (fig. 50a). the general lza field, nevertheless, transgresses towards the equilibrium cotectic proportions and the overlying gabbro zones. however, these do not reach the cotectic proportions predicted https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 82 of 139 geusbulletin.org for basaltic systems on the ol-di join (l(pl,cpx,ol)) of fig. 47). the lzb can be interpreted as a mixture of c. 60% olivine gabbro and c. 40% trapped melt (fig. 50b). the lzc and mz are made up of gabbros with relatively similar clinopyroxene content and trapped melt of 15–20% (fig. 50c). for uza, we estimated c. 10% trapped melt (fig. 50d). these values of trapped melt, although also displaying an upward decrease, are significantly higher than those estimated by tegner et al. (2009) and in this study (see sections 5.11.4 and 5.12). 5.7.7 field of liquid immiscibility the modal variation in the uz displays a marked increase in olivine near the base of uzb, followed by decreasing olivine and a concurrent increase in augite upward in uzb (fig. 51). we consider the possibility that this marked increase in olivine near the base of uzb is a consequence of liquid immiscibility by comparing our observations to the experimental immiscibility results of charlier & grove (2012). fig. 52a (projection from pl) and fig. 52b (projection from di) compare the experimental melt just prior to immiscibility (bc, binodal cusp) and the two conjugate immiscible melt compositions (“fe” and “si” in fig. 52; compare to fig. 47a). the modelled liquid compositions at the base of uzb correspond to the cusp-melt compositions of charlier & grove (2012) at the unset of immiscibility. there is also as a good correspondence between the si-rich conjugate melt (“si” in fig. 52) and the various granophyre compositions observed in the skaergaard intrusion (“bt”, “t” and “g” in fig. 52). figure 52c compares the uz gabbros to the experimental immiscibility results. the first order observation is that the uzb and uzc fields intersect at the location of uza and branch out from here towards either normative q (or granophyre) for uzc or towards ol (or troctolite) for uzb. this suggests that uza is made up of gabbros formed prior to the onset of immiscibility. after immiscibility reaches the liquidus at the base of uzb, partial separation of the two conjugate melts results in uzb gabbros affected by enrichment in the basic-melt component and the uzc gabbros affected by enrichment in the silicic melt. using the experimental determination of the two conjugate-melt compositions (charlier & grove 2012), uzc is made of a mixture of gabbro and up to 30% granophyre – not so different from the maximum uzc content of 20% granophyre detected in this study (see section 5.4). the silicic conjugate melt and the granophyre are probably located near pseudo-invariant crystallisation (grove et al. 1982) and are saturated in pyroxene, albite (an~0 and not an30 as in the host gabbros), quartz and probably orthoclase consistent with the uzc near the sh (figs 13–15). the uzb gabbros can be interpreted as mixtures of gabbros and basic conjugate melt (fig. 52c) in relatively constant proportions of normative di/q (or clinopyroxene/granophyre) and highly variable olivine (fig. 51). the content of the conjugate basic melt varies from 6% to 20% and is positively correlated with diopside (or clinopyroxene) content. the strong enrichment in normative olivine is consistent with the observed modal content (fig. 15). in effect, this ultimately results in a magma chamber originating from enrichment in the basic conjugate melt at the base of uzb and enrichment in the silicic conjugate melt at the top of the chamber and in the ubs. with further cooling, the remaining magma will eventually exit the immiscibility field and proceed to solidify upward from the base of the uzb and downward from the ubs, terminating at the sh. the basic conjugate melt resulting from liquid immiscibility falls within the olivine field and can be predicted after termination of immiscibility to crystallise this phase, initially without large amounts of plagioclase and augite (fig. 52c) and thus resulting in an enrichment of olivine at the base of the uzb (fig. 51). with increasing crystallisation, the modal content of olivine decreases, and augite and plagioclase increases, and eventually augite is restored on the liquidus. the silicic conjugate plots near quartz saturation or would eventually end in a crystallising quartz together with albitic plagioclase. however, in the projections both conjugate melts fall close to a pigeonite field and could eventually be expected to become saturated in pigeonite (e.g. grove et al. 1982). despite this, orthopyroxene crystallisation appears to have been terminated in upper uza and does not appear in uzb or uzc. the reason for this is not clear but could be related to a thermal divide on the 1000 1200 1400 1600 1800 2000 2200 0 1 2 3 modal olivine/clinopyroxene (wt%) st ra ti gr ap hi c he ig ht (m ) sh uzc uzb uza mz fig. 51 variation in the ratio between modal olivine and clinopyroxene (molar weight fractions) as a function of stratigraphic height (m) in the uz. subzone divisions for mz, uza, uzb and uzc and sh are shown (abbreviations in fig. 2). https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 83 of 139 geusbulletin.org reaction curve between pigeonite and orthopyroxene as suggested by grove et al. (1983). it could also be an artifact of not incorporating ferric iron in the immiscibility melt compositions, or a limitation in our understanding of the low-temperature phase equilibria of silicic melts. an interpretation of the temperature of onset of liquid immiscibility on the skaergaard liquid lines of descent is shown in fig. 53 as a function of the melt structure measured by non-bridging oxygens per tetrahedral-coordinated cations (nbo/t). the modelled skaergaard liquids (tables 21, 22) reach into the two-liquid domain, although without any experimentally-derived sign of immiscibility (thy et al. 2006). this penetration occurs at a higher temperature (c. 1060°c) than expected from the delineation of the two-liquid field of charlier & grove (2012; c. 1000–1020°c), as well as for other experimental studies of immiscibility in natural basalts (dixon & rutherford 1979; philpotts & doyle 1983). veksler et al. (2007, 2008) have, nevertheless, proposed liquid unmixing at temperatures of 1110–1120°c based on centrifugation experiments on a composition believed to be representative for the mz of the skaergaard, although such high temperatures have been widely challenged (philpotts 2008; morse 2008b; charlier & grove 2012) and are not di (pl) liquid immiscibility(mol%) ol q uzb si-enrichment sife di q ol opx pl qol (di) opx fe si l (pl,cpx,ol) liquid immiscibility b (mol%) m g bt t di q pl opx ol uzb si-enrichment opx m t conjugate melts: fe: basic si: silicic b: binodal cusp a l (pl,cpx,ol) pl b b conjugate melts: fe: basic si: silicic b: binodal cusp uza c uzc di (pl) liquid immiscibility uzb (mol%) ol q uza granophyre aug gab bro 10% 20% 30% si-enrichment ave. max. min. sife liquid immiscibility (charlier & grove 2012) troctolite uzb, uzc b conjugate melts: fe: basic si: silicic b: binodal cusp di q ol opx pl bt g liquid immiscibility (charlier & grove 2012) liquid immiscibility (charlier & grove 2012) m g bt t granophyre tindenfell basistoppen melanogranophyre fig. 52 interpretation of liquid immiscibility. projections from plagioclase (pl) and diopside (di) of the quaternary ol–q–pl–di on the triangular di–ol–q and pl–ol–q diagrams (see quaternary perspective diagrams, inset) in mol%. projection method in fig. 45. a: experimental result of charlier & grove (2012) projected from pl. b: experimental result of charlier & grove (2012) projected from di. c: interpretation of uz gabbros based on the results of charlier & grove (2012). the binodal cusp melt is denoted by ‘b’, which resolved to the fe-rich and si-rich conjugate melts with cooling. the granophyre melts of table 18 are given by square symbols marked by ‘g’ (granophyre), ‘m’ (melanogranophyre), ‘t’ (tindenfell) and ‘bt’ (basistoppen) in panels a, b. the additions of granophyre are marked on the uza-granophyre join (10%, 20%, 30% in panel c). solid curves: liquid lines of descent. max.: maximum. ave.: average. min.: minimum. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 84 of 139 geusbulletin.org supported in this study. nevertheless, intermediate temperatures in the range of 1050–1060°c may be possible, although not yet demonstrated experimentally (charlier & grove 2012; honour et al. 2019a). honour et al. (2019b) examined immiscibility in tholeiitic glasses and found an apex temperature for initiation of immiscibility at about 1125–1115°c (kilauea iki and laki lavas), based on the mgo content of the glass. such various estimates of the onset of immiscibility are thought to be controlled by compositional and melt structure in the bulk melts with compositions ranging from alkalic, to transitional or olivine tholeiitic and quartz tholeiitic. the onset of immiscibility in the skaergaard intrusion is interpreted to have occurred at the uza–uzb boundary at c. 1060°c and to have been saturated in plagioclase (an40), olivine (fo39), clinopyroxene (mg# 50), feti oxides and possibly apatite at that point. the two modelling options suggest a 65% sio2 content in melt and 14% feo* (for si enrichment) or 54% sio2 and 26% feo* (for fe enrichment). the former shows some similarities with the results of charlier & grove (2012) for their starting composition for the sept iles sill, suggesting 60% sio2 and 10% feo*. their corresponding minerals are plagioclase (an37), olivine (fo43), clinopyroxene (mg# 54), feti oxides, apatite and pigeonite. although there are some common mineralogical and compositional characteristics, the skaergaard immiscibility appears to have occurred for melt compositions significantly less evolved and at a higher temperature than those for the sept iles sill. these differences are not explained but may be related to the relatively alkalic nature of the sept iles sill as revealed by its augite compositions. philpotts (1979) found that the temperature of onset of immiscibility would rise from 1020°c to 1050°c as a result of increasing the log fo2 from –13 to –9. or in other words: with increasing fugacity, liquid immiscibility will occur at earlier stages of crystallisation. the present modelling of the liquid lines of descent suggests an log fo2 at the onset of immiscibility of between –10.2 and –12.6, respectively for fe enrichment and si enrichment (fig. 53). thus, a lower fo2 cannot have contributed to the unusual high temperature at immiscibility in the skaergaard intrusion compared to the experiment at –9.0 fmq of charlier & grove (2012). it is also plausible that the difference in melt compositions between the skaergaard liquids and those included in the experiments of charlier & grove (2012) may contribute to some of the anomalies encountered for the olivine-hypersthene normative-parental skaergaard melt. charlier & grove (2012) demonstrated that immiscibility is possible for large sio2 and feo* melt variations, including in effect both liquid lines of melt model considered here. they also demonstrated that na2o and non-quaternary oxides (k2o, p2o5 and tio2) would promote immiscibility at a lower feo* enrichment. fractional crystallisation of parental skaergaard melts, following augite and subsequently feti oxides crystallisation, would result in systematic increases in sio2 accompanied by na2o, k2o and p2o5 enrichments (with variable feo*). it is possible that such variations in fractional crystallisation may promote immiscibility at higher temperatures than implied by the equilibrium experimental determination of charlier & grove (2012), even though bulk-melt compositions may have been similar. the melt structure described by the nbo/t ratio suggests that the skaergaard intrusion reached the field of liquid immiscibility in polymerised melts with nbo/t values of 0.32 and 0.55, respectively, for the si-enriched and fe-enriched liquid lines of descent (fig. 53). this is significantly higher levels of polymerisation than for the initial skaergaard melt at 0.80 and is the result of substantial augite and feti oxide crystallisation. using the lever rule, the relative proportions of the two conjugate immiscible melts can be estimated at 84–72% silicic and 16–28% basic melts, with the range given for the two liquid decent models, respectively. if the nbo/t ratios were extrapolated to the bimodal surface, without considering the effects of fractional crystallisation, si enrichment fe enrichment 920 960 1000 1040 1080 1120 1160 0 0.4 0.8 1.2 1.6 2 lld two-liquids �eld sk sk m g t bt charlier & grove (2012) nbo/t t (° c ) uzb uza nbo/t: pl: 0.0; cpx: 2.0; ol: 4.0 –9.2 (log fo2)–13.5 –9.0 (log fo2) one-liquid �eld –10.2 –12.6 granophyre transgressive sills basistoppen jakobsen et al. (2011) melanogranophyre g t bt m sk fig. 53 interpretation of the temperature (t; °c) of onset of liquid immiscibility as a function of the degree of polymerisation measured by non-bridging oxygens divided by temperature (nbo/t). the evolution paths for the two liquid line of descent calculations are from tables 21 and 22. italic numbers are modelled log fo2 (thy et al. 2006). granophyre compositions are based on those in table 18. the two analyses marked ‘sk’ and connected by a horizontal green dashed line are the conjugate melt estimated for the skaergaard intrusion by jakobsen et al. (2011, arbitrarily plotted at 1040°c). the lower boundary (grey dashed curve) between the ‘one liquid’ and ‘two-liquids’ fields are from charlier & grove (2012). the upper boundary (black dashed curve) is predicted from the skaergaard intrusion raised to the location of the uza–uzb boundary. pl: plagioclase. cpx: clinopyroxene. ol: olivine. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 85 of 139 geusbulletin.org we would obtain ratios of about 1:1 – more akin to equilibrium crystallisation than fractional crystallisation as assumed in the calculations of the uz liquid lines of descent (see sections 5.11.9 and 5.13). the liquid lines can be predicted to have penetrated from the immiscibility field into either a low-temperature single silicic melt field, by reaching a low-temperature subliquidus surface (charlier & grove 2012), or from a breakdown of immiscibility due to removal of silicic melt from the system. however, the relevant phase relations that would allow firm predictions are not known (fig. 53). for an immiscibility field nested on a cotectic (as in the case of the skaergaard magma) to leave a detectable imprint on the liquid descent and solid fractionation products, disturbances of the immiscible melts are required such as gravitative separation or coalescence of one or both of the cognate melts. without disturbances and with cooling and crystallisation, the immiscible field will disappear and leave no imprint on magmatic evolution (roedder 1978, 1979), except perhaps the droplets trapped in zoned primocrysts and interstitial crystallisation products (jakobsen et al. 2005, 2011). the olivine-rich base of the uzb is such a disturbance in the immiscibility processes. 5.7.8 apatite crystallisation primocrystic apatite appears suddenly and at relatively high abundance reaching 8% at the boundary between uza and uzb. the apatite abundance positively correlates with olivine (fig. 54) and suggests that apatite, like the concurrent olivine base, may also be an effect of a liquid immiscibility event. watson (1979) showed apatite saturation to be a function of temperature and melt composition (particularly sio2 and p2o5) in basaltic to andesitic melts. tollari et al. (2006) further showed that cao had a large effect on apatite saturation, while varying iron and oxygen fugacities had little effect. thus, the p2o5 content for apatite saturation decreases with increasing sio2 and cao (watson 1979; green & watson 1982; tollari et al. 2006). consequently, apatite may be stable in very felsic melts like granophyres (watson & capobianco 1981). our forward modelling (table 21) suggests crystallisation along the si-enriched liquid line of descent using the saturation model of tollari et al. (2006) when the melt reaches c. 2.0 wt% p2o5, similar to the prediction of tegner et al. (2009, c. 1.7% p2o5), when the uza–uzb boundary is taken to be 1060°c (fig. 55). in contrast, apatite saturation is never reached along the fe-enriched liquid line of descent using the same saturation model as tollari et al. (2006). this means that apatite would reach stability at the uza–uzb boundary for a si-enrichment trend only and would remain stable at low temperatures in silicic derived melt. using the partitioning of p2o5 between the two conjugate melts determined by watson (1976; db/s = 10), the p2o5 content of the basic conjugate melt would, because of the high d value, have had nearly the same p2o5 content (c. 1.8 wt%). this means that most of p2o5 would go into the basic conjugate melt and that the presence of apatite in the skaergaard intrusion could be controlled by liquid immiscibility and not crystal fractionation. it is thus not possible to determine conclusively whether apatite appearance at the uza–uzb boundary is a result of protracted fractionation or the result of immiscibility and the accumulation of basic conjugate melt at the boundary. both scenarios would be able to explain the upward decreasing modal content of apatite. this would be for fractional crystallisation due to decreasing p2o5 and for liquid immiscibility due to upward increasing granophyre the positive correlation between modal apatite and the ‘excess’ olivine at the olivine (wt%) 0 5 10 15 20 25 30 0 2 4 6 8 a pa ti te (w t% ) fig. 54 modal content of apatite (wt%) as a function of modal olivine (wt%). curve is visual estimate and excludes outliers. 1050 11401090 1100108010701060 113011201110 0 1 2 8 7 6 5 4 3 9 t (°c) a pa ti te s at ur at io n at m el t p 2o 5 (w t% ) lld si-enrichment apatite saturation tollari et al. (2006) uzauzb no apatite apatite apatite saturation fig. 55 the p2o5 concentrations (wt%) of melt saturated in apatite as a function of temperature (t; °c), calculated using the equation of tollari et al. (2006) and the forward modelled liquid line of descent (lld si-enrichment; table 21). the two curves only intersect at the uzb–uza boundary (indicated by grey arrow), and thus apatite may thus first have become a liquidus phase at the onset of liquid immiscibility in the skaergaard intrusion. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 86 of 139 geusbulletin.org base of uzb is nevertheless seen as a strong indication that the appearance of apatite, like olivine, is related to the onset of liquid immiscibility. 5.7.9 bulk-rock compositions and cotectic relations it is conspicuous that the majority (lzc–uza) of the skaergaard gabbros (figs 46, 47) cluster around the experimental determination of the cotectic for basaltic melts saturated in plagioclase, olivine and clinopyroxene (thy & lofgren 1992, 1994). this would imply that the gabbros formed from the accumulation of three primocryst silicate phases (plus feti oxides and apatite) and trapped melt, without recording excessive phase separation either from crystallisation or gravitative sorting. this is clearly not the case since there is abundant evidence that extreme primocryst sorting or separation is related to extensive layering in the outcrops of the skaergaard intrusion (e.g. wager & deer 1939; irvine 1987; maaløe 1987; naslund et al. 1987; conrad & naslund 1989). the good correspondence with expected phase boundaries can be explained simply by the field-work approach of targeting and sampling average gabbros, in addition to the three-sample averages used in the projections, which have been reasonably successful in reducing, but not eliminating, the scatter in the diagrams. the reasonable correspondence can also be explained where the gabbros are made of mostly included elements and that excluded elements have little effect on bulk-gabbro compositions and the phase diagrams as constructed. sampling of average gabbros is commonly used for understanding the evolution of layered intrusions as pioneered by wager (1960). 5.8 iron-magnesium exchange equilibria equilibria between coexisting pyroxenes and olivine can be evaluated by the exchange reaction coefficient for fe and mg between olivine and pyroxenes, expressed by the binary kd calculated as: 0 200 400 600 800 1000 1200 1400 1600 1800 2000 2200 0 0.5 1 1.5 2 2.5 3 3.5 kd (ol/cpx) kd (opx/cpx)kd (opx/ol) st ra ti gr ap hi c he ig ht (m ) kd = (fe/mg)m1/(fe/mg)m2 uzc uzb uza mz lzc lzb lza fig. 56 calculated partitioning kd of fe and mg for three different pairs of pyroxenes and olivine (tables 3, 5 and 6) as a function of stratigraphic height (m) in the ls. calculated using eq. (4) with all iron as fe2+. 0 1 2 3 0 1 2 0 1 2 3 0 1 2 0 1 2 3 a cb (fe/mg)cpx (fe/mg)ol (fe/mg)cpx olivine orthopyroxene orthopyroxene measured observed k d = 1 k d = 1 k d = 1 k d = 1 .4 k d = 0.7 y = 2.2 x1.4 (r2 = 0.93) (f e/ m g) o l o r ( fe /m g) o px ? fig. 57 co-variation of fe/mg ratios for observed orthopyroxenes and olivine (table 3, 5 and 6) together with the results for one-atmospheric melting experiments (measured; thy et al. 2006). all iron has been calculated as fe2+. a: (fe/mg)ol (olivine) vs. (fe/mg)opx (clinopyroxene) pairs. the fe/mg ratio can be fitted to the power equation shown, suggesting kd values between 1 and 2.5. b: (fe/mg)opx (orthopyroxene) vs. (fe/mg)ol pairs. the data can be fitted to a linear equation giving kd = 0.7. c: (fe/mg)opx vs. (fe/mg)opx pairs. the data can be fitted to a linear equation giving kd = 1.4, although a power equation is a possibility (curve marked ‘?’). https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 87 of 139 geusbulletin.org kd = (fe/mg)m1 / (fe/mg)m2 (4) where fe and mg are mole fractions, m1 and m2 are coexisting mafic silicates (olivine, clinopyroxene or orthopyroxene), and fe is calculated as total iron. fig. 56 shows kd as a function of stratigraphic height (or generally decreasing temperature) in the ls for three mineral pairs (see also figs 31, 32). for all pairs, kd markedly changes at the uza–uzb boundary, where liquid immiscibility is suggested to commence (see sections 5.7.7 and 5.12). until the mz, the olivine-clinopyroxene pair shows a constant kd of 1.73 ± 0.11 (n = 14), after which kd systematically increases to c. 3 until reaching uzc, where the observed values are highly variable (fig. 56). the marked increase in kd with decreasing t was also observed by loucks (1996), reaching well above 2 for temperatures below 1000°c. these values are comparable to the experimental determinations of thy et al. (2006), suggesting kd = 1.41 ± 0.29 (n = 40) for temperatures above 1080°c (fig. 57). there is thus an indication of exponential increase in kd with increasing fe and decreasing temperature (fig. 57a). the orthopyroxene-olivine pair shows a constant kd = 0.73 ± 0.08 (n = 24) until orthopyroxene ceases to crystallise in the uzb (fig. 57b). this kd is similar to the experimental kd = 0.74 ± 0.03 (n = 4) obtained by thy et al. (2006) for the skaergaard dykes. these latter results, however, contrast with experimental results by medaris (1969) and observations for the bushveld complex and stillwater complex (atkins 1969; raedeke & mccallum 1984; barnes & naldrett 1986) typically indicating values of c. 0.9–1.0 (fig. 57b). this difference may be an effect of the relatively evolved iron-rich skaergaard melts compared to the more mafic compositions of the other layered intrusions. the important observation is that there is a good correspondence between the observed kd and the experimentally derived kd for the skaergaard intrusion (fig. 57b). the orthopyroxene-clinopyroxene pair suggests constant partition until the uza with an average kd of 1.33 ± 0.16 (n = 24), followed by a slight increase in uza (fig. 57c). the kd for pyroxenes from layered intrusions varies between 1.40 and 1.28 (brown et al. 1957; mccallum 1968; atkins 1969; saxena 1973; fleet 1974; barnes & naldrett 1986), similar to the values for skaergaard kd observed here. these values are higher than most of the experimental determinations of thy et al. (2006) for skaergaard dykes (kd = 1.12 ± 0.10 (n = 6) as well as the test for equilibrium of putirka (2008, p. 95; kd = 1.09 ± 0.14; fig. 57c). again, there is some indication of an exponential increase in kd with increasing fe and decreasing temperature, similar to, but less prominent than, in the relationship between the olivine and clinopyroxene pair. the reason for the moderate, but distinct, difference between observed and measured experimental results is not clear, although the results may mostly be within the experimental and analytical uncertainty. the graphical pyroxene thermometer of lindsley & andersen (1983) indicates consistent temperatures for the average clinopyroxene compositions (fig. 58) suggesting that the compositions obtained in this study, using a broad-beam electron microprobe technique, have reasonably captured the solidus compositions for clinopyroxene. this is also supported by the good match 800 9001000 700 800 900 600 500 1000 1100 1200 1300 1200 1100 pig aug opx bustamite fsen di he augite pigeonite bustamite measured fig. 58 pyroxene quaternary diagram with phase equilibria (thin black lines) and temperature contours (solid and dashed grey lines) after lindsley & andersen (1983). the pyroxene compositions of tables 3 and 5 calculated according to lindsley & andersen (1983) reduce the pyroxenes to their quadrilateral en–di–he–fs components and take all iron to be calculated as fe2+. black dashed tie-lines are shown for representative skaergaard pyroxenes. orange dashed tie-line is for the experimentally determined, average coexisting pyroxenes. bustamite (inverted ferrobustamite) of uzc from table 7 is shown for comparison. di: diopside. he: hedebergite. en: enstatite. fs: ferrosilite. opx: orthopyroxene. pi: pigeonite. aug: augite. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 88 of 139 geusbulletin.org to values obtained by previous studies (brown et al. 1957; brown & vincent 1963) based on mineral separates (fig. 19) as well as the general lack of strong subsolidus and hydrothermal imprints on the average clinopyroxene compositions (nwe 1975, 1976; manning & bird 1986). in contrast, orthopyroxene often suggests much lower equilibrium temperature than for clinopyroxene by up to 100ºc (fig. 58). the exception is at the mz–uza transition, where coexisting pyroxenes appear to approach equilibrium and better match solidus relations (compare figs 19 and 58). this suggests that the majority of the skaergaard orthopyroxene compositions may record a strong component of subsolidus equilibration caused by exsolution and grain coarsening, consistent with the phase equilibria considerations (see section 5.7.5). the effect of solidus to subsolidus equilibration for the pyroxenes is expected to increase the fe/mg ratio of orthopyroxene and decrease the ratio for clinopyroxene. this leads to an anticlockwise rotation of the tie lines for coexisting pyroxenes in the quaternary (see coleman 1978) and eventual widening of the solvus (lindsley 1983). a comparison between the coexisting pyroxenes obtained in melting experiments (small filled circles in fig. 58) and skaergaard gabbros (large circles in fig. 58) shows the expected widening of the pyroxene gap for the cumulates with decreasing temperature. also, despite nearly identical pigeonite-saturation temperature for the gabbro primocrysts (c. 1090°c) and experiments (c. 1084°c), higher pyroxene fe/mg ratios are suggested by the experiments than are observed for primocrysts. further, there is a clear mismatch between the observations and the contours drawn by lindsley & andersen (1983). it is often assumed that the fe/mg ratios for equilibrated pyroxenes define a kd approaching 1.0 (coleman 1978; putirka 2008). in fig. 58, this would mean that the tie lines for high and low ca pyroxenes would extrapolate to a point close to the wollastonite apex. however, for the skaergaard intrusion, this does not appear to be the case. rather, the tie lines for the experimental pyroxene pairs extrapolate to a ca/mg ratio of c. 2.7 similar to the orientation of tie lines for gabbro rocks with similar pyroxene mg numbers. this suggests that values of kd for the coexisting pyroxenes in highly evolved basaltic melts are dependent on temperature and composition. in summary, despite these differences, the consistency of fe-mg exchange in the experiments on skaergaard-like liquids and gabbroic rocks suggest that the pyroxenes in the gabbros are well equilibrated, albeit at subsolidus conditions. 5.9 magmatic volatiles although the initial skaergaard magma is considered to have been virtually dry, evidence of volatile components (h,cl,f,s) is captured by apatite, sulphides and hydrous silicates in the gabbros, particularly after advanced degrees of crystallisation. assuming an initial volatile content equivalent to basalts from the galapagos spreading centre (h2o = 0.09%, cl = 0.08%, f = 0.02%, s = 0.12%; byers et al. 1984; michael & chase 1987), and assuming perfect fractional crystallisation, we can estimate the concentrations of volatile components at the beginning of uzb and at the onset of liquid immiscibility. using a trapped-liquid content of 2.5% (tegner et al. 2009), the volatile content can be calculated as h2o = 0.71 wt%, cl = 0.24 wt%, f = 0.08 wt% and s = 0.56 wt%. these values are taken to represent the end-member estimates at the uza–uzb boundary, as well as for an equivalent interstitial-melt evolution below this boundary. the estimates for h2o are well below saturation for basaltic and ferrobasaltic magmas and the precipitation of amphibole and biotite (burnham 1975; berndt et al. 2005; shishkina et al. 2010). however, the modelled h2o content (0.36–0.71) may have been sufficiently high to result in detectable effects on the mineralogy and crystallisation temperatures in at least the mz to uza (berndt et al. 2005; botcharnikov et al. 2008). the experimental calibration of almeev et al. (2007, 2012) predicts that in the upper part of uza, the liquidus depression for olivine may have reached 15–30°c, depending on the trapped-melt model, and 35–60°c for plagioclase at the top of the uza, relative to the dry liquidus. the experiments of almeev et al. (2007, 2012) were performed on typical morb composition and an applied pressure of 200 mpa. the relatively strong depression suggested for the plagioclase liquidus would markedly have modified the melt evolution and equilibrium compositions of plagioclase in the upper parts of the intrusion, while similar effects on olivine saturation would not have been as marked as for plagioclase. a small h2o content may also have raised the temperature of the onset of liquid immiscibility (cf. lester et al. 2013) and may partially explain the inferred high temperature of immiscibility for the skaergaard (see section 5.7.7). however, the effects of wet (or damp) crystallisation have so far eluded experimental and forward modelling of the skaergaard intrusion, which have exclusively been modelled assuming dry conditions (toplis & carroll 1996; thy et al. 2006). the reason for this is the lack of applicable experimental calibrations and information on melt composition. biotite, intermediate between phlogopite and annite (fig. 24), occurs as a minor interstitial and late-crystallising mineral together with ilmenite and apatite (fig.  6). this hydrous silicate was considered by nash (1976) to have been secondary, subsolidus reaction products, involving a residual magmatic fluid. hydrothermal, mineralised fracture veins are also abundant in the intrusion (bird et al. 1986, 1988), some of which contain biotite of https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 89 of 139 geusbulletin.org a similar composition to those of this study. nash (1976) interpreted the fluorine and h2o partitioning between apatite and biotite to suggest that biotite had lost fluorine to a circulating meteoric fluid, while apatite had retained the magmatic fluorine content. however, using the analytical results of nash (1976) as well as sallet’s (2000) solution to the partitioning of fluorine between coexisting biotite and apatite, an approximate equilibrium temperature of 920°c is suggested. this may be a reasonable estimate of near-solidus temperature, above most of the subsolidus temperatures, based on the feti oxide thermometry. however, it ignores the substantial amount of chlorine in apatite (table 10). the appearance of magmatic biotite, confined to the lz and into the base of the mz, is attributed to a channelled influx of water into parts of the early mush in approximate contact to the host, metamorphic basement. such an influx could have resulted in sufficiently raised local h2o concentrations to saturate biotite without reaching h2o saturation and fluid development. pegmatitic gabbros in the lz contain trapped fluid inclusions with h2o and ch4, suggesting that the final evolved interstitial melt and the sh alike may have coexisted with a hydrous fluid similar to that trapped in the minerals at t < 770°c (larsen et al. 1992; sonnenthal 1992; larsen & brooks 1994). the monovalent anions in apatite suggest a twostep boiling point. first as a loss of chlorine in the uzb, where apatite becomes an early-forming mineral, and liquid immiscibility reaches the liquidus, and again in uzc, when apatite essentially becomes dry by losing oh (fig. 24b). mcbirney (1995, 2002) and sonnenthal (1992) observed an increase in the f/cl ratio of apatite from the beginning of the upper part of mz and culminating at the sh (or slightly above), which they attributed to the development and removal of chlorine with a hydrous fluid vapor (cf. doherty et al. 2013). they also suggested that the upward concentration of excluded trace elements near the top of the intrusion was caused by a gradual change in the partition coefficient in response to the exsolution of a volatile fluid. this formation of a cl-rich and volatile-rich fluid may have influenced the development of the mineralisation in the upper part of the mz (boudreau & mccallum 1992). detailed petrographic and geochemical studies have partly confirmed this expectation and argue that the formation of au-rich sulphides in the topmost part of the mineralisation was associated with a cl-bearing aqueous fluid (godel et al. 2014; nielsen et al. 2015; rudashevsky et al. 2023). sulphide liquid immiscibility occurred when the liquid sulphur content reached 0.26 wt% in the upper part of the mz according to andersen (2006), compared to our estimate of 0.22–0.16 wt% at approximately the same level. this saturation resulted in strong concentrations of the au and pges in a narrow zone in the upper part of the mz (triple group; andersen et al. 1998), although this interpretation is controversial (bird et al. 1991; andersen et al. 1998; andersen 2006; godel et al. 2014; keays & tegner 2015; nielsen et al. 2015, 2019b; holwell et al. 2016; pedersen et al. 2020). the presence and the effects of volatile components on the fractionation history of the skaergaard intrusion are poorly understood, although evidence points towards the precipitation of hydrous silicates and to the separation of fluid vapours from the melt in the upper part of the intrusion. future studies should focus on volatile components as well as experiments and modelling of their effects on the liquid lines of descent and trace element partitioning to expand our understanding of the potential importance of volatiles, particularly for the late stages of solidification. 5.10 kinematic and thermal constraints an evaluation of the dynamic behaviour of melt and primocrysts is indicative for understanding processes forming the mush and its final solidification mechanisms. the density relations of the mush system can be calculated for both primocrysts and coexisting melt using partial molar volumes (bottinga & weill 1970; bottinga et al. 1982; lange & carmichael 1987, 1990; knoche et al. 1995). the densities of the primocrysts and crystal network are calculated here using the thermal expansion equations of niu & batiza (1991a, b), taking into consideration the mineral composition, liquidus temperature and observed mineral modes. the single mineral and crystal network-density calculations are shown in fig. 59a. the calculated density of the crystal network can be compared to the specific gabbro density measured relative to water using archimedes’ principle (fig. 59a). the measured densities, although matching the observed densities, are slightly lower, presumably because the former is affected by cracking at ambient temperatures (tegner et al. 2009). the composition of the coexisting melt is not directly known. however, density calculations using the two end-member modelled liquid lines of descent are reasonable predictions of the possible ranges in density (tables 21, 22). the calculation results are compared in figs 59a and 59b, showing the two liquid lines of descent and the effects of a small amount of water amounting to an initial 0.10 wt% (see discussion on magmatic volatiles in section 5.9). the primocrysts group into low-density plagioclase (c. 2.6–2.7 g·cm−3), medium-density mafic silicate minerals (3.4–4.0 g·cm−3) and high-density feti oxide minerals (3.7–4.9 g·cm−3; fig. 59a). of the primocrysts, https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 90 of 139 geusbulletin.org only olivine suggests a strong increase in density with stratigraphy, reaching a difference of 0.7 g·cm−3 from the base to the top in the intrusion, thus recording the upward increasing fayalite. the density of the coexisting melt is predicted to be very similar to, or narrowly parallels, that of plagioclase (fig. 59a). the difference between plagioclase primocryst and melt densities varies for the si-enriched liquid line of descent from –0.15 to +0.15 g·cm−3 with the neutral crossover in the middle of uza, a potential buoyancy of plagioclase in the lower part of the intrusion and an accumulation in the upper part of the intrusion (fig. 59a). for the fe-enriched liquid line of descent, the liquid is systematically higher by 0.13–0.17 g·cm−3 throughout the intrusion, suggesting plagioclase buoyancy throughout the ls. the effect of water as well as a high trapped-liquid content are minimal until about the uzb (fig. 59b), where liquid immiscibility appears on the liquidus and the density calculations that use the liquid lines of descent based on fractional crystallisation may have less practical relevance. the viscosity of a magmatic melt is a measure of its internal resistance to flow, the tendency for primocryst migration in the magma and for melt migration in the mush network. dynamic viscosity is measured as the ratio of the shear stress to the time rate of shear strain, is dependent on temperature and composition, and in particular, on the silica melt content or melt polymerisation. the viscosity has been directly determined experimentally in complex and simplified systems using rotational or falling sphere viscometers (e.g. ryan & blevins 1987). such experimental results have been used by bottinga & weill (1972), shaw (1972) and giordano et al. (2008) to formulate equations for the temperature and compositional dependency of viscosity for complex silicate melts. the results of the viscosity calculations for the skaergaard liquid lines of descent, using the procedure of giordano et al. (2008), are shown in fig. 59c. for 0 80 160 240 320 400 stoke’s law calculation for 2-mm particles dry si-enrichment liquid line of descent d 6.55.54.53.52.51.5 modelled melts si fe fe si c 2.82.62.42.22.0 modelled melts si fe si fe b uzc uzb uza mz lzc lzb lza hz ap 3.72 mt 4.95 il 4.74 fe ti o xi de 5.04.54.03.53.02.52.0 2200 2000 1800 1600 1400 1200 1000 800 600 400 200 0 –200 –400 pl c px c px ol ol m el ts fe si fe si pl a settling velocity (m⋅y–1)viscosity log η (pa·s)density (g⋅cm–3)density (g⋅cm–3) st ra ti gr ap hi c he ig ht (m ) conjugate melts si (wet) si (dry) fe (wet) fe (dry) plagioclase pyroxene olivine feti oxides modelled melts: gabbro (measured) gabbro (calculated) conjugate melts conjugate melts c on ju ga te m el ts wet drydry wet feti oxides gabbro ol pl cpx fig. 59 density (g·cm–3) and viscosity (pa·s) as a function of stratigraphic height (m) in the ls. a: calculated density of primocrysts (pl: plagioclase. cpx: clinopyroxene. ol: olivine), gabbro and feti oxides. pl, cpx and ol are calculated using the temperatureand compositional-dependent equations of niu & batiza (1991a, b). pressure effects are minor within the range assumed for the skaergaard intrusion (60–130 mpa) and are not considered here. the density of low-ca pyroxene has been calculated together with high-ca pyroxene (marked as cpx). densities of apatite (ap), magnetite (mt) and ilmenite (il) are also shown and assumed to be typical values. the densities of gabbro and feti oxide are based on calculated modal contents (table 2). measured densities of gabbro are shown for comparison and use archimedes’ principle. the range in melt variation between the two, end member modelled compositions are marked as si and fe (see supplementary tables s1 and s2, respectively, and panel b for details). the conjugate melts (si and fe) are from charlier & grove (2012) and are given as averages, with 1σ sd. b: a focussed view of the calculated densities of the two, end-member liquid lines of descent (si and fe) from panel a. grey shading shows the variation in melt for the dry liquidus. the wet liquidus is calculated for an initial h2o content of 0.1 wt. conjugate melts in a. c: calculated dynamic viscosities (pa·s) of the modelled melts. the fields for the dry and wet liquidi are shown for the two, end-member liquid lines of descent. conjugate melts in a. d: the settling viscosity (m·y-1) calculated from stokes’ law for 2 mm diameter particles or clusters. calculated from the dry si-enrichment liquid line of descent (table s1). https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 91 of 139 geusbulletin.org the si-enrichment trend with increasing silica content, the viscosity shows an increase from 1.8 to 6.2 – on a logarithmic scale in fig. 59c – but a more restricted variation for the fe-enrichment trend of 1.8 to 3.4 – again on a logarithmic scale. for the former, this is equivalent to 90 pa·s at the base of lza, to 71 000 pa·s at the base of uzb. multiply by 10 to convert these viscosities to the traditionally used poise unit based on g·cm–1·s–1. water systematically lowers the viscosity by 9–18% on the logarithmic scale. stokes equation relates the contrast between viscosity and density to the settling velocity of primocrysts in a stagnant or slowly convecting magma as: v = (2g . δρ . r2) / (9η) (5) where v is the terminal settling velocity (m·s–1), g is the gravitational acceleration (9.8 m·s–2), δρ, either positive or negative, and is the density contrast between the primocryst and the melt (kg·m–3), r is the radius of a spherical primocryst (m), and η is the dynamic viscosity (pa·s). this equation allows us to calculate the settling or flotation rate and thus the separation of primocrysts in melts of known or assumed densities and viscosities. using stokes equation, eq. (5), the settling or flotation rate for plagioclase grains (of a constant 1 mm radius) in a stagnant anhydrous melt of the lza is about 30 m·y–1 and markedly decreased upward to approach zero at the base of the uza or slightly negative values above (fig. 59d). the other main silicate primocrysts (olivine and pyroxene), assuming a similar size, will settle with a rate of 98–132 m·y–1 and again will become stagnant at uzb. the gabbro rate is calculated assuming a clustered settling that represents the observed average gabbro mode and its variation. all these calculated settling rates are normalised to a unifying particle of 2 mm in diameter. for the si-enriched liquid line of descent, the gabbro settling rate is 43 m·y–1 in lza reaching near zero in uzb. using the fe-enrichment descent will slightly increase this rate to 51 m·y–1 in lza and, because of the more restricted silica enrichment in the upper part of the intrusion, will stabilise in uza only at a rate of <10 m·y–1. water approximately doubles the settling velocity in the lz, caused mainly by the break-up of silicate melt polymerisation and thus the reduction in viscosity. the total settling times for clustered gabbro accumulation, integrated from the velocity data shown in fig. 59d, suggest fast accumulation in the lower part of the intrusion (lz–uza) of 307 years (an average of 3.7 m·y–1) and for the whole intrusion of 4130 years (an average of m·y–1). these accumulation times apply to the si-enrichment liquid line of descent. as expected, they only increase if the fe-enrichment liquid line of descent is considered (an average of 13 m·y–1) and similarly for wet conditions (18–34 m·y–1). these settling durations are several orders of magnitudes faster than rates estimated for the skaergaard intrusion from thermal heat-transfer modelling (wager & brown 1967, p. 210; irvine 1970b; norton & taylor 1979). the thermal models assume heat loss mainly through the roof and adiabatic convection in the melt and incorporate temperature contrasts between magma and the host, specific heat content and density, thermal conductivities and latent heat of crystallisation. the modelling of norton & taylor (1979, fig. 10) suggested a duration of 130 000 years to reach the solidus temperature (or f = 0) and 80 000 years to reach uzb or a temperature of about 1060°c (f = 0.25). other estimates, orders of magnitude lower, have been proposed by wager & brown (1967, p. 210) who used a constant accumulation rate of 0.20 m·y–1 based on the modelling of the stillwater complex by hess (1960) and thus suggested 12 000 years for the entire ls and an inferred 8000 years to reach the uzb. the “basic model”, incorporating the development of an ubs, as formulated by irvine (1970b, fig. 10), suggests an upward accumulation rate of approximately 11 m·y–1 and thus a total time of 40 000 years for the entire ls and 22 000 years to reach the uzb – significantly less than the estimates of norton & taylor (1979). the apparent low estimate by wager & brown (1967) is due to not incorporating a ubs that significantly reduces the heat loss from the roof of the intrusion (irvine 1970b; irvine 1974). similar cooling estimates for the bushveld complex, to about the same minimum temperature as used in irvine’s calculations (1970b), suggest accumulation rates of 0.04–0.05 m·y–1 (cawthorn & walraven 1998; zeh et al. 2015). the use of stokes equation assumes spherical particles in an infinite, non-convecting media without interaction with other particles to calculate the steadystate terminal velocities. thus, the modelling probably has limited applications to the skaergaard crystallising chamber. specifically, the calculations are normalised to unified-sized particles and non-spherical particle shapes, and particle interactions in the multi-saturated melt are certainly plausible. further, the gabbro clustering in question probably cannot be reduced to mm-sized spherical composite particles but would more likely involve solid porous-particle settling in a melt. such factors may thus greatly reduce the settling velocity by as much as 100 times (irvine 1974) than the ideal conditions shown in fig. 59d. furthermore, adiabatic convection in the main magma chamber may cause delay or retention of crystals (marsh & maxey 1985; marsh 1988; martin & nokes 1989) – particularly in high viscosity melts like the upper part of the skaergaard intrusion, https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 92 of 139 geusbulletin.org but less so in the lower parts where settling velocities may have been higher. despite such uncertainties and limitations in the application of stokes equation to magma chambers, the normalised settling rates of fig. 59d suggest that crystal settling might have operated in the lower part (i.e. lz) of the intrusion and is unlikely to have had any effect in the upper parts (i.e. mz and uz). there is also no strong indication of plagioclase flotation, although stagnation may have been in operation (cf. maaløe 1984). the possible effects of liquid immiscibility can be evaluated using the experimental results of charlier & grove (2012). calculated on a dry basis, the two conjugate immiscible melts have densities with differences of about 0.50 g·cm–3 (fig. 59b), probably sufficient to result in liquidus separation of these two melts. further, the contrast in the dry viscosity for the two conjugate liquids is high with 270 pa·s for the fe melt and 54 000 pa·s for si melt (fig. 59c). the separation velocity for the two melts is thus well over 100 m·y–1, based on a settling velocity of 112 m·y–1 for the si-rich cognate melt and 0.6 m·y–1 for the fe-rich melt. this suggests the cognate immiscible melts might have been able to efficiently separate with the si-rich melt floating upward and the fe-rich melt sinking or remaining stagnant at the base of uzb. the surface tension (gas-melt) of the two individual conjugate immiscible melts can be estimated by the addition of the partial molar tensions at 1500°c using the procedure of mills (1986). the results adjusted to a temperature of 1060°c, as estimated for immiscibility in the fractionating skaergaard melt, by using the dependence determined by walker & mullins (1981), is for the fe-rich melt 265 ± 2 dyn·cm–1 (equivalent to mn·m–1) and for the si-rich melt 173 ± 9 dyn·cm-1. the procedure of kucuk et al. (1999), adjusted as above from a temperature of 1400°c, suggests 346 ± 1 dyn·cm-1 and 301 ± 5 dyn·cm-1, respectively. these are within the range of the experimental results obtained on a variety of basaltic and andesitic melts by walker & mullins (1981), by veksler et al. (2010) for the simplified k2o-feo-fe2o3al2o3-sio2 system, and by murase & mcbirney (1973) on basalts to rhyolites as read from their graph at an extrapolated temperature of 1060°c. the results indicate similar surface tensions with the fe-rich immiscible melt slightly higher than for the si-rich melt. based on these individual estimated values, we can calculate a dry interfacial tension (melt-melt) of 2–10 dyn·cm-1 between the two conjugate melts by using the solution of girifalco & good (1957) and the simplification suggested by veksler et al. (2010). this is a low interfacial tension, but not very different to those directly measured by veksler et al. (2010) at 8–16 dyn·cm-1 for corresponding immiscible melts in the k2o-feo-fe2o3-al2o3-sio2 system. veksler et al. (2010) further speculated that the interfacial tension in natural basaltic melts would be 2 or 3 times lower than observed in their simplified system, or in other words corresponding to what is estimated here. such a low interfacial tension would have resulted in enhanced nucleation in the case of fe-rich melt droplets in a continuous si-rich melt (charlier & grove 2012; honour et al. 2019a). another effect would be to slow coarsening by ripening and coalescence of the droplets, which would result in a protracted stability of the emulsion and hence restrict early gravitative separation of the immiscible melts (veksler & charlier 2015). the latter appears to contrast with the evidence for large-scale gravitative separation of the emulsion in the uzb suggested for the skaergaard intrusion (see section 5.12). this may, however, be an apparent rather than a real conflict, since the time duration for uzb–uzc to solidify has variably been estimated to thousands of years (e.g. norton & taylor 1979). such long timespans may have allowed near complete coarsening of melt droplets (james 1975) and thus allow sufficient time for breakdown of any initial emulsions and the effective separation of the two immiscible melts. although interfacial tension may be relevant for understanding the fluid-dynamic behaviour of the immiscible melts in the residual melt chamber, it is less significant when considering what is happening in the trapped melt in the mush zone where solid-melt interactions with primocrysts (or wetting properties) may in part control the relative migrations of the two immiscible melts in highly heterogeneous, but poorly understood mush pathways (honour 2019; honour et al. 2019b; holness et al. 2020). 5.11 forward modelling of trace elements the vast amount of field data, petrographic data and compositional data for the skaergaard lends itself to more rigorous testing of liquid line of descent models. we have already discussed the modelling of the major elements (see sections 5.5–5.8). here we continue by examining the trace elements to evaluate various reference fractionation models. the elemental concentrations of the gabbros (tables  1, 16) represent a mixture of primocrysts and trapped interstitial (or mesostasis) material (wager 1963; henderson 1970, 1975). the calculated mineral modes in table 2 are based on these bulk-gabbro compositions, which include only primocrysts and are without a component of trapped melt. the results may thus only represent a first order approximation to the primocryst mineral modes. this may, nevertheless, be a https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 93 of 139 geusbulletin.org reasonable approximation given that the trapped melt typically appears to be restricted (tegner et al. 2009) and may have crystallised without significant migration and melt fractionation. the excluded trace elements are confined to late-formed minerals but may also be present in small amounts in the primocrysts. an example, which is often used to calculate the trapped melt content, is rb that may reside in minute and late-crystallised orthoclase, but also enters plagioclase in small amounts. this may result in a bulk-partitioning coefficient for rb in the mz of around 0.14 (see section 5.11.3). our modelling is based on the bulk-gabbro compositions and experimentally known partitioning. first, we evaluate the mineral-melt partitioning (di), then the bulk partitioning (di *), and then the effects of trapped-melt component (ftm). we then show that in some cases reference fractionation models fail to fully account for the observations and we are consequently steered in new directions, sometimes still without satisfactorily solving the problem at hand. we begin with some basic constraints on element partitioning followed by forwarding modelling of the trace element variations. 5.11.1 calculation methods the trace element partition (di) between coexisting solid mineral and liquid is defined on a weight basis as: di = ci m / ci l (6) where di is the nernst partition coefficient and ci is the concentration of an element i, respectively, in mineral (m) and liquid (l). the bulk partition coefficient (di *) between the solid and the liquid for an element i for j = 1 to n can be defined as: di * = ∑mi j × di j (7) or by inserting eq. (6) into eq. (7), for j = 1 to n, we obtain: di * = ∑mi j · ci j / ci l (8) where mi j is the modal weight fraction of a specific mineral j and a specific element i, di j is the elemental partition coefficient of eq. (6) for the same mineral j, and n is the total number of minerals (or components) in the gabbro rock. the modal weight fractions (mi) vary individually from 0 to 1, and total 1. the di * can be estimated for many elements because the modal solid mineral fractions are known (table 2; fig. 15) and the individual elemental partition coefficients can be reasonably estimated using existing experimental data (e.g. nielsen 1992). equations (7) and (8) can be expanded to n + 1 by adding an interstitial trapped-melt component with di = 1, treating the trapped-melt component like a mineral and assuming little or no migration of melt. calculated this way, d* represents the crystal mush, incorporating a trapped-melt fraction that theoretically can vary from zero to 1. elements with d* > 1 are defined as included in the bulk mush (compatible), while elements with d* < 1 are defined as excluded from the mush zone (incompatible) but included in the residual melt. the residual (or inverted) liquid concentration (ci l) can be obtained by inversion because the di * and gabbro mush concentration are known. thus, the equilibrium concentration of a trace element in the liquid can be predicted using the bulk partition coefficient from eq. (8) and the gabbro concentration, resulting in: ci l / ci mush = 1 / di * (9) where ci mush is the measured concentration of the element i in the gabbro. only for di * = 1 is ci l equal to ci mush; describing perhaps a hypothetical end-condition where the mush is made up of 100% ‘trapped’ melt. perfect fractional crystallisation can, if f is known, be calculated stepwise as ci l / ci o using the batch equilibrium-crystallisation equation, ci l / ci o = 1 / (di * (1 – f) + f) (10) or by the rayleigh distillation equation, ci l / ci o = f (d i *– 1) (11) where ci l is the resultant concentration and ci o is the previous melt composition for individual steps, di * is constant for an individual step, and f is the ratio between f (here the liquid-fraction remaining) for the resultant and f for the previous melt segment. f for perfect fractional crystallisation can be estimated from the stratigraphic height using the method of tegner et al. (2009) in eq. (1). if the initial concentration of an element i is known or can be constrained otherwise, ci l can thus be calculated for each step from eqs (10) or (11). equilibrium crystallisation can be approximated using eq. (9), assuming that ci mush is constant for each step and equal to the initial melt composition, and not variable as for perfect fractional crystallisation. the equilibrium variation is constrained by the bulk partition coefficient, the modal make-up of the gabbros, and ultimately the temperature and the initial concentration. equilibrium crystallisation is specifically of interest for evaluating the effects of imperfect fractional crystallisation that can be calculated as a mixture of fractional and equilibrium crystallisation. used this way, ‘imperfect’ does not refer to disequilibrium distributions of the trace elements. the fraction of trapped melt (ftm) in the gabbros for each step can be estimated using an excluded element (i) as: https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 94 of 139 geusbulletin.org ftm = ci mush / ci l (12) where ci l is the predicted concentration of element i in the liquid using eq. (10) or (11) and ci mush is the measured gabbro or mush concentration of the same element. ftm approaches an ideal value as di * → 0. using eq. (11) or (12) based on the assumption of perfect rayleigh fractionation, tegner et al. (2009) calculated the skaergaard trapped-melt fractions using the gabbro concentrations for rb or u throughout the ls. similar calculations are iteratively done here assuming stepwise di *-values initially without a priori incorporating a trapped-melt content. in situ crystallisation can, following langmuir (1989), be modelled as ci l / ci o = f d (d* – 1) / (d* (1 – d) + d) (13) where c, f, and d* are defined previously and d is the fraction of melt in the boundary zone returned to the main chamber. the boundary zone where crystallisation occurs can be viewed as the sum of the fractions of solid, trapped liquid, and the fraction of the returned liquid, such that d is equal to one minus the combined solid fraction and ftm. eq. (13) is similar to the equation by langmuir (1989) for the case without trapped melt and is applicable here because d* incorporates the trapped melt component and, argues langmuir (1989), approximates the partitioning between the returned melt and the original melt. the limiting values of d = 1 correspond to perfect fractional crystallisation, while d = 0 equates to equilibrium crystallisation. the trapped melt fraction is known and the solid fraction of the crystallising boundary zone can be bracketed, or estimated, to be between 0 and 1 with a reasonable constant average of 0.5, assuming a linear variation from base to top as suggested by langmuir (1989). 5.11.2 mineral-liquid partition coefficients (d) the partitioning of the elements sc, v, cr, co, ni, zn, rb, sr, y, zr, nb, ba, la, yb and u has been evaluated for olivine, plagioclase, pyroxenes, feti oxides and apatite, using the experimentally determined mineral-liquid partition coefficients (d) reported by nielsen (1992) as a fortran code taking into account the dependency on temperature, melt composition, and some aspects of mineral composition. the lattice-compensated strain model of blundy & wood (1994), as formulated by bédard (2005, 2006, 2007, 2014), was used to calculate the partition for the ree together with y and sc. temperature can be estimated from the an content of plagioclase using the method of thy et al. (2009a, 2013) and the mineral compositions in tables 3–5. the major element melt compositions are from the forward modelling in tables  21 and 22, calculated to represent the two contrasting views of fe depletion and si enrichment (hunter & sparks 1987; toplis & carroll 1996) and prolonged fe enrichment with restricted si enrichment (mcbirney & naslund 1990; tegner 1997), both past feti oxide crystallisation and both extrapolated to f = 0. the use of the two contrasting liquid lines of descent thus results in two sets of partitioning coefficients. a summary of d values for minerals and subzones is shown in table 23 where values overlap below the mz, as expected. one major limitation in these calculations is the uncertainty in the modelled liquid lines of descent, but there also exist uncertainties in the partitioning relations for silica melts in the models of nielsen (1992) and bédard (2014) that could become problematic towards the end stages of evolution for strong si-enrichment models. the main conclusion is that the trace element partitioning depends on temperature as well as melt and mineral compositions. the individual partition coefficients for the prolonged fe-enrichment model (lower end of the ranges in table 23) are similar to those obtained by mcbirney (1998), who only assumed a melt composition characterised by increasing fe enrichment beyond feti oxides (mcbirney & naslund 1990). as in our modelling, mcbirney (1998) also used the source code of nielsen (1992) for his calculations. the transition metals (sc, v, cr, co, ni, and zn) are partitioned principally into magnetite and ilmenite, sc, cr and co also into the pyroxenes, and co and ni also into olivine (table 23). the partition coefficients increase for most of the elements with height in the stratigraphy, caused by decreasing temperature, melt evolution and mineral variables as also demonstrated by nielsen et al. (1992, 1994) and nielsen & beard (2000). on the other hand, the differences in the calculated partition coefficients for cr, ni, sc and v (table 23) for the two liquid descent models can only be caused by differences in the melt compositions, since both temperature and mineral compositions are identical for the two models. in contrast to the other transition metals, ni is principally incorporated into olivine (dni = 35 at the mz level), to a lesser extent into clinopyroxene (dni = 17), and orthopyroxene (dni = 8; table 23). it is well established that dni for olivine is dependent on melt composition to the extent that a marked increase with decreasing forsterite content, decreasing temperature and melt fraction of sio2 is plausible (hart & davis 1978; kinzler et al. 1990; bédard 2005; li et al. 2003; li & ripley 2010). thus, the increase and divergence in dni for olivine in the uz can be attributed to variation and differences in melt sio2, again considering that temperature and olivine compositions are identical for the two liquid line of descent models. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 95 of 139 geusbulletin.org ta bl e 23 s um m ar y of s el ec te d si ng le -e le m en t a nd b ul kel em en t p ar tit io n co effi ci en ts sc v cr co n i zn rb sr y zr n b ba la yb u o liv in e u zc 1. 4– 3. 3 0. 06 –0 .1 5 0. 58 –1 .4 2. 0– 4. 7 69 –1 64 1. 1– 2. 7 0. 22 –0 .6 3 0. 42 –1 .1 u zb 1. 4– 3. 3 0. 06 –0 .1 1 0. 50 –0 .8 8 1. 9– 3. 3 57 –1 02 1. 1– 1. 9 0. 23 –0 .6 4 0. 43 –1 .1 u za 1. 1– 1. 8 0. 05 –0 .0 7 0. 35 –0 .4 8 1. 5– 2. 0 40 –5 4 0. 85 –1 .1 0. 17 –0 .3 3 0. 33 –0 .5 9 m z 0. 60 –0 .6 5 0. 04 –0 .0 5 0. 29 –0 .3 3 1. 3– 1. 5 33 –3 7 0. 74 –0 .8 6 0. 06 –0 .0 6 0. 12 –0 .1 4 lz c 0. 45 –0 .6 5 0. 04 –0 .0 4 0. 26 –0 .2 8 1. 2– 1. 3 28 –3 1 0. 70 –0 .7 6 0. 04 –0 .0 4 0. 09 –0 .1 0 lz b 0. 33 –0 .4 9 0. 04 –0 .0 4 0. 22 –0 .2 4 1. 2– 1. 2 24 –2 6 0. 65 –0 .6 9 0. 03 –0 .0 3 0. 06 –0 .0 7 lz a 0. 27 –0 .2 9 0. 04 –0 .0 4 0. 21 –0 .2 1 1. 1– 1. 1 22 –2 3 0. 63 –0 .6 5 0. 02 –0 .0 2 0. 04 –0 .0 5 cl in op yr ox en e u zc 40 –5 6 0. 60 –1 .4 2 10 3– 24 7 2. 7– 6. 5 49 –1 17 0. 66 –1 .6 2. 0– 4. 7 2. 0– 2. 7 0. 42 –4 .3 0. 07 –0 .1 7 0. 39 –0 .5 8 1. 7– 2. 2 u zb 15 –2 9 0. 49 –0 .9 9 66 –1 19 2. 6– 4. 5 36 –6 4 0. 62 –1 .1 1. 4– 2. 5 1. 3– 1. 8 0. 31 –4 .7 0. 07 –0 .1 2 0. 21 .0 .3 4 1. 1– 1. 5 u za 8. 0– 13 0. 45 –0 .6 1 33 –4 5 2. 1– 2. 8 22 –2 9 0. 49 –0 .6 6 0. 79 –1 .1 0. 9– 1. 1 0. 13 –1 .4 0. 06 –0 .0 7 0. 14 –0 .1 8 0. 86 –1 .0 m z 5. 6– 6. 0 0. 40 –0 .4 6 21 –2 5 1. 8– 2. 1 16 –1 8 0. 43 –0 .5 0 0. 56 –0 .6 4 0. 73 –0 .7 6 0. 10 –0 .1 6 0. 05 –0 .0 6 0. 10 –0 .1 0 0. 68 –0 .7 1 lz c 5. 0– 5. 2 0. 37 –0 .4 2 16 –1 8 1. 7– 1. 9 13 –1 4 0. 41 –0 .4 5 0. 45 –0 .4 9 0. 65 –0 .6 7 0. 10 –0 .1 2 0. 04 –0 .0 5 0. 09 –0 .0 9 0. 62 –0 .6 3 lz b 4. 4– 4. 5 0. 35 –0 .3 7 11 –1 2 1. 6– 1. 7 10 –1 1 0. 38 –0 .4 0 0. 34 –0 .3 6 0. 59 –0 .6 0 0. 11 –0 .0 9 0. 04 –0 .0 5 0. 08 –0 .0 8 0. 57 –0 .5 7 lz a 4. 0– 4. 1 0. 34 –0 .3 5 8. 9– 8. 9 1. 5– 1. 6 8. 3– 8. 5 0. 37 –0 .3 8 0. 27 –0 .2 8 0. 55 –0 .5 5 0. 24 –0 .0 9 0. 04 –0 .0 4 0. 07 –0 .0 7 0. 53 –0 .5 3 o rt ho py ro xe ne u zc u zb u za 4. 3– 8. 9 0. 30 –0 .4 0 7. 7– 10 1. 8– 2. 4 11 –1 4 0. 42 –0 .6 2 0. 06 –0 .1 5 0. 57 –0 .7 9 m z 1. 9– 2. 2 0. 26 –0 .3 0 6. 2– 7. 1 1. 6– 1. 8 7. 9– 8. 9 0. 24 –0 .2 6 0. 02 –0 .0 2 0. 36 –0 .3 8 lz c 1. 4– 1. 5 0. 24 –0 .2 7 5. 4– 6. 0 1. 5– 1. 6 6. 4– 7. 0 0. 17 –0 .1 9 0. 01 –0 .0 1 0. 26 –0 .2 9 lz b 1. 0– 1. 1 0. 23 –0 .2 5 4. 7– 5. 0 1. 4– 1. 5 5. 0– 5. 3 0. 11 –0 .1 3 0. 00 –0 .0 1 0. 19 –0 .2 1 lz a 0. 78 –0 .8 3 0. 22 –0 .2 3 4. 2– 4. 3 1. 4– 1. 4 4. 1– 4. 2 0. 09 –0 .1 0 0. 00 –0 .0 0 0. 15 –0 .1 6 pl ag io cl as e u zc 0. 09 –0 .1 1 0. 07 –0 .1 7 0. 16 –0 .3 8 1. 5– 1. 0 0. 09 –0 .0 9 0. 20 –0 .0 1 0. 18 –0 .1 8 0. 07 –0 .0 7 u zb 0. 09 –0 .1 2 0. 07 –0 .1 2 0. 15 –0 .2 6 0. 24 –0 .3 4 0. 07 –0 .0 7 0. 03 –0 .0 5 0. 16 –0 .1 6 0. 05 –0 .0 5 u za 0. 08 –0 .1 0 0. 06 –0 .0 7 0. 12 –0 .1 6 0. 43 –0 .8 7 0. 06 –0 .0 6 0. 08 –0 .1 5 0. 14 –0 .1 4 0. 04 –0 .0 4 m z 0. 04 –0 .0 5 0. 05 –0 .0 6 0. 11 –0 .1 2 0. 74 –0 .7 9 0. 05 –0 .0 5 0. 14 –0 .1 5 0. 13 –0 .1 3 0. 04 –0 .0 4 lz c 0. 02 –0 .0 3 0. 04 –0 .0 5 0. 10 –0 .1 1 0. 74 –0 .7 0 0. 05 –0 .0 5 0. 15 –0 .1 4 0. 12 –0 .1 2 0. 03 –0 .0 3 lz b 0. 01 –0 .0 1 0. 04 –0 .0 5 0. 09 –0 .1 0 0. 91 –0 .6 7 0. 04 –0 .0 4 0. 20 –0 .1 5 0. 11 –0 .1 1 0. 03 –0 .0 3 lz a 0. 00 –0 –0 0 0. 04 –0 .0 4 0. 09 –0 .0 9 1. 5– 0. 76 0. 04 –0 .0 4 0. 37 –0 .1 8 0. 10 –0 .1 0 0. 02 –0 .0 3 (c on tin ue d) https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 96 of 139 geusbulletin.org ta bl e 23 (c on tin ue d) s um m ar y of s el ec te d si ng le -e le m en t a nd b ul kel em en t p ar tit io n co effi ci en ts sc v cr co n i zn rb sr y zr n b ba la yb u m ag ne ti te u zc 4. 2– 4. 5 44 –1 05 75 3– 17 90 13 –3 0 21 –5 1 19 –4 6 0. 07 –0 .1 9 0. 09 –0 .2 2 u zb 4. 2– 4. 5 41 –7 3 62 6– 11 11 12 –2 1 19 –3 4 18 –3 2 0. 07 –0 .1 9 0. 09 –0 .2 2 u za 3. 9– 4. 2 33 –4 5 43 5– 58 5 9. 6– 13 15 –2 0 14 –1 9 0. 07 –0 .1 9 0. 09 –0 .2 2 m z 3. 0– 3. 2 29 –3 4 34 7– 39 9 8. 2– 9. 7 12 –1 4 13 –1 5 0. 04 –0 .0 9 0. 06 –0 .1 1 lz c 2. 3– 2. 5 27 –3 0 30 4– 33 5 7. 9– 8. 7 11 –1 3 12 –1 3 0. 03 –0 .0 5 0. 04 –0 .0 6 lz b 1. 5– 1. 7 26 –2 7 26 1– 27 7 7. 4– 7. 9 10 –1 1 11 –1 2 0. 04 –0 .0 4 0. 05 –0 .0 5 lz a 1. 0– 1. 1 25 –2 6 23 3– 24 1 7. 1– 7. 4 9. 7– 10 11 –1 1 0. 14 –0 .0 4 0. 16 –0 .0 5 ilm en it e u zc 16 –1 8 20 –4 7 44 –1 04 2. 4– 5. 8 18 –4 3 0. 50 –0 .5 6 11 –2 6 27 –6 3 0. 39 –0 .5 4 0. 51 –0 .5 6 u zb 16 –1 8 19 –3 3 41 –7 3 2. 3– 4. 0 17 –3 0 0. 50 –0 .5 6 10 –1 8 25 –4 4 0. 40 –0 .5 4 0. 52 –0 .5 6 u za 14 –1 5 15 –2 0 33 –4 5 1. 8– 2. 5 14 –1 8 0. 46 –0 .4 9 8. 2– 11 20 –2 7 0. 32 –0 .4 0 0. 48 –0 .5 1 m z 9. 2– 10 13 –1 5 29 –3 4 1. 6– 1. 9 12 –1 4 0. 31 –0 .3 3 7. 2– 8. 3 18 –2 0 0. 11 –0 .1 4 0. 36 –0 .3 8 lz c 6. 2– 7. 1 12 –1 3 27 –3 0 1. 5– 1. 7 11 –1 2 0. 22 –0 .2 5 6. 7– 7. 4 17 –1 8 0. 04 –0 .0 6 0. 28 –0 .3 0 lz b 3. 1– 3. 9 12 –1 2 26 –2 7 1. 4– 1. 5 10 –1 1 0. 12 –0 .1 4 6. 3– 6. 7 16 –1 7 0. 01 –0 .0 2 0. 17 –0 .2 0 lz a 1. 5– 2. 0 11 –1 1 25 –2 6 1. 4– 1. 4 10 –1 0 0. 06 –0 .0 8 6. 1– 6. 3 15 –1 6 0. 00 –0 .0 0 0. 10 –0 .1 3 a pa ti te u zc 0. 22 –0 .5 9 0. 81 –1 .9 2. 1– 5. 0 14 –5 3 0. 51 –1 .2 9. 1– 34 7. 2– 37 0. 80 –1 .9 u zb 0. 22 –0 .4 2 0. 76 –1 .3 2. 0– 3. 5 13 –3 4 0. 48 –0 .8 4 8. 4– 24 6. 6– 22 0. 75 –1 .3 u za 0. 18 –0 .2 3 0. 61 –0 .8 2 1. 6– 2. 1 10 –1 7 0. 38 –0 .5 2 6. 8– 12 5. 3– 9. 6 0. 60 –0 .8 1 m z 0. 08 –0 .0 9 0. 53 –0 .6 1 1. 4– 1. 6 8. 8– 11 0. 34 –0 .3 9 6. 1– 7. 2 4. 4– 5. 6 0. 53 –0 .6 1 lz c 0. 04 –0 .0 5 0. 50 –0 .5 5 1. 3– 1. 4 8. 0– 8. 7 0. 32 –0 .3 5 5. 7– 6. 1 3. 9– 4. 4 0. 49 –0 .5 4 lz b 0. 03 –0 .0 3 0. 47 –0 .5 0 1. 2– 1. 3 7. 2– 7. 3 0. 30 –0 .3 1 5. 3– 5. 4 3. 5– 3. 6 0. 46 –0 .4 9 lz a 0. 03 –0 .0 3 0. 45 –0 .4 7 1. 2– 1. 2 6. 8– 7. 0 0. 29 –0 .3 0 5. 1– 5. 3 3. 3– 3. 4 0. 45 –0 .4 6 bu lk p ar ti ti on c oe ffi ci en ts b as ed o n m in er al m od es in t ab le 2 a nd t ra pp ed m el t by t eg ne r et a l. (2 00 9) u zc 15 –2 1 1. 2– 2. 8 38 –9 1 1. 5– 3. 5 31 –7 6 0. 46 –1 .1 0. 09 –0 .1 6 1. 3– 1. 8 1. 0– 1. 9 1. 2– 3. 3 1. 3– 3. 2 0. 49 –0 .4 7 0. 36 –0 .8 1 0. 93 –1 .7 0. 19 –0 .2 3 u zb 5. 0– 8. 6 2. 1– 3. 6 29 –5 0 1. 4– 2. 4 22 –3 9 0. 73 –1 .2 0. 09 –0 .1 4 0. 50 –0 .8 5 0. 90 –1 .8 0. 76 –2 .2 1. 6– 2. 8 0. 05 –0 .0 8 0. 48 –1 .1 0. 67 –1 .4 0. 06 –0 .1 0 u za 4. 0– 5. 8 2. 8– 3. 7 31 –4 1 1. 4– 1. 8 31 –4 1 0. 87 –1 .1 0. 12 –0 .1 4 0. 49 –0 .7 6 0. 46 –0 .5 5 0. 76 –1 .4 1. 7– 2. 3 0. 12 –0 .1 5 0. 21 –0 .2 4 0. 43 –0 .5 2 0. 08 –0 .0 9 m z 3. 2– 3. 5 3. 4– 3. 9 28 –3 2 1. 4– 1. 6 11 –1 5 0. 89 –1 .0 0. 14 0. 52 –0 .6 0 0. 39 –0 .4 1 1. 1– 1. 2 2. 4– 2. 8 0. 16 –0 .1 7 0. 19 –0 .2 0 0. 38 –0 .4 0 0. 11 lz c 2. 6– 2. 8 3. 3– 3. 7 25 –2 7 1. 3– 1. 5 7. 4– 8. 1 0. 89 –1 .0 0. 16 0. 48 –0 .4 9 0. 38 –0 .3 9 1. 1– 1. 2 2. 4– 2. 6 0. 13 0. 21 0. 38 –0 .3 9 0. 14 lz b 1. 3– 1. 4 0. 77 –0 .8 0 6. 2– 5. 6 0. 82 –0 .8 6 4. 1– 4. 4 0. 4 0. 24 –0 .2 5 0. 66 –0 .5 6 0. 37 –0 .3 8 0. 34 0. 47 –0 .4 9 0. 29 –0 .2 7 0. 28 0. 36 –0 .3 7 0. 21 lz a 0. 86 –0 .8 1 0. 77 –0 .8 9 3. 8– 3. 9 0. 80 –0 .5 5 3. 2– 3. 9 0. 5 0. 38 0. 92 –0 .6 4 0. 44 0. 47 –0 .4 5 0. 57 0. 48 –0 .4 1 0. 40 0. 43 –0 .4 4 0. 35 ca lc ul at ed u si ng n ie ls en ’s (1 99 2) a nd b ed ar d’ s (2 00 5, 2 00 6, 2 00 7, 2 01 4) a lg or ith m s an d m el t c om po si tio ns a nd te m pe ra tu re s de sc ri be d in th e te xt . t he r an ge s gi ve n re pr es en t t he e xt re m e liq ui d lin es o f d es ce nt a s pr op os ed fo r th e sk ae rg aa ar d in tr us io n w ith f een ri ch m en t fi rs t, fo llo w ed b y si -e nr ic hm en t. u nd er lin ed p ar tit io n co effi ci en ts a re fo r in cl ud ed e le m en ts . o rt ho py ro xe ne p ar tit io ni ng c oe ffi ci en t w as n ot c al cu la te d fo r u zb ,c . s in gl e va lu es r ec or d si m ila ri ty fo r th e tw o m el t m od el s. o rt ho cl as e an d zi rc on p ar tit io ni ng in u zc w er e in cl ud ed a s co ns ta nt s ba se d on g eo ch em ic al r ef er en ce m od el (g er m ; h ttp s: // ea rt hr ef .o rg /g er m /). https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ https://earthref.org/germ/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 97 of 139 geusbulletin.org the alkali elements rb and ba are principally partitioned into the melt with only small amounts going into plagioclase, resulting in drb = 0.11 and dba = 0.14 at the mz level. for plagioclase, drb increases and diverges markedly through uz, as do the d values for many of the transition elements. experimentally determined drb and dba values increase as a function of decreasing an of plagioclase and mgo and sio2 melt contents (bédard 2006), supporting the general findings of nielsen (1992). on the other hand, sr is predicted to be partitioned principally at the level of mz, if present, into apatite (dsr = 1.5) and to a lesser extend into plagioclase (dsr = 0.76) and clinopyroxene (dsr = 0.60). the calculated irregular variation of dsr for plagioclase as a function of stratigraphic height contrasts to the summary of bédard (2006) who predicted an increase in dsr with decreasing an content of plagioclase and increasing sio2 of melt. strontium, however, is known to be included in apatite (prowatke & klemme 2006; watson & green 1981) and to a lesser extend in clinopyroxene (bédard 2014), consistent with the modelling results. the high-field strength elements zr and nb partition into ilmenite (dzr = 7.5 and dnb = 19) and for zr also to a much lesser extent into apatite (dzr = 0.36) and clinopyroxene (dzr = 0.13), all reported at the mz level. the modelled values for ilmenite partitioning are high compared to experimental determinations of dzr and dnb by klemme et al. (2006) and parker et al. (2011) at, or just below, unity for synthetic ti-rich silicate systems. the ree together with y partition into apatite with dy = 9–11 and dree = 4–7 at the mz level and diverge through the uz like other trace elements related to differences in the increase in sio2 consistent with the observations of watson & green (1981) and prowatke & klemme (2006). ree and y are also moderately partitioned into clinopyroxene (dla = 0.10, dyb = 0.70, and dy = 0.70, all at the mz level). bédard (2014) reviewed the existing experimental information on trace element partitioning into clinopyroxene. he observed that the partitioning is dependent on temperature and melt composition and that many elements in felsic and mafic melts formed coherently, slightly increasing d-trends with decreasing temperature and other compositional variables, reflecting advanced fractionation, although the most evolved felsic melts were rarely found to be collinear with intermediate and mafic melts. 5.11.3 bulk partition coefficients (d*) the calculations of the di * values are summarised in table 23 for the individual zones and subzones. used for the calculation of di * are the observed average modes (table 2), the estimated trapped-melt content after tegner et al. (2009), and the calculated di values summarised in table 23. in addition, the partition in uzc into orthoclase for rb, sr and ba was based on the geochemical earth reference model database (germ 2016), and the partition for zircon in uzc was likewise included, based on a small amount (0.02 wt%; wager & brown 1967). because of its low concentrations of trace elements, quartz was not included in the calculations. other minor mineral phases are not considered (e.g. sulphides and alloys); this later omission means that cu could not be modelled. fig. 60 shows the variation in di * for selected elements (sc, v, rb, ba, yb and u) as a function of stratigraphic height. the results allow a preliminary definition of the skaergaard trace elements into either excluded (di * < 1.0) or included (di * > 1.0) elements – the latter is identified by underlined values in table 23. the majority of the transition elements are consistently included, while of the excluded elements, only rb, ba, and u are truly excluded (<0.15) due to the presence of apatite, suggesting that these elements are reasonable candidates for trapped-melt estimates (e.g. tegner et al. 2009). other minor mineral phases are not considered (e.g. sulphides and alloys); the later omission means that cu could not be modelled. by definition, values of di * depend on the modal content. because the trapped melt content is included in the calculations, di * will approach unity with increasing trapped-melt content at the base of the ls (tegner et al. 2009); di * increases for excluded elements and decreases for included elements due to the content of trapped melt (table 23). the two modelled liquid lines of descent shown in fig. 60, from f = 0.76 to f = 0, are initiated at identical values f = 1, below the base of the stratigraphic column because of the presumably similar melt compositions (h = 365 m). the di * values may thus vary as a function of melt composition and stratigraphy but are constrained by the initial value f = 1. the differences seen in the di *-values, if any, are mostly a function of differences in the melt compositions and modal mineralogy, not in temperature, since that is identical for the two melt models. the calculated di * values reflect the large scatter in the bulk-gabbro modal make-up, even after being reduced by the stepwise average compositions (table  13), although the scatter is somewhat reduced compared to those observed by mcbirney (1998, 2002). common for many of the predominantly included elements is an upward transition from excluded to included status (tables 23). for v, sc, co, zn and nb, the di * values increase with modal feti oxide minerals in the lzc to mz. of these elements, only sc shows a positive correlation with modal clinopyroxene, consistent with the high di value for this phase (table 23). surprisingly, there is no clear positive correlation between cr content and modal pyroxene or feti oxide, even though this would https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 98 of 139 geusbulletin.org have been expected from the di values. only the sr content shows positive correlation with modal plagioclase as reflected by di just below unity and likewise only ni correlates positively with modal olivine, but not with modal pyroxene as again expected from the di values. of the excluded elements, only rb, ba and u are consistently excluded throughout the ls, and their d* values are consequently strongly affected by the trapped-melt component in the lza and lzb (fig. 60; table 23). ree and y are excluded elements in the lower part of the intrusion, but transition towards an included status in the uzb with the appearance of apatite. the effect of the melt composition on di * is first manifested in the mz where the two liquid lines of descent models start to diverge, resulting in differences in the uz by a magnitude of approximately two. this is mainly attributed to the differences in the terminal si content (56% versus 74% sio2) resulting in increased di * values for the latter liquid line of descent. one significant limitation to this modelling is the assumption we made at the outset that f = 0 at the top of uzc. while this constraint was imposed by tegner & cawthorn (2010) in computing their liquid lines of descent, it may not strictly be the case. for example, the forward models of andersen (2006), nielsen et al. (2009) and hunter & sparks (1987), respectively, assumed 2%, 5% and 18–25% liquid remaining after crystallisation of uzc. the di * values may thus vary as a function of stratigraphy, but the differences are mostly functions of melt compositions and modal mineralogy, and only slightly due to temperature. 5.11.4 trapped melt content the ftm was modelled by tegner et al. (2009) using excluded elements (p, rb and u) and di values iteratively adjusted for a trapped-melt content assuming that di * are equal to the trapped-melt fraction (cf. fig. 60, rb). the results showed a marked decrease in ftm from a lza lzb lzc mz uza uzb uzc st ra ti gr ap hi c he ig ht (m ) d* melt concentration (ppm) gabbro concentration (ppm) d* melt concentration (ppm) gabbro concentration (ppm) sc d* gabbro concentration (ppm) d* gabbros melt concentration (ppm) d* d* gabbro concentration (ppm) melt concentration (ppm) d* d* gabbro concentration (ppm) gabbros melt concentration (ppm) yb d* d* gabbros melt concentration (ppm) gabbro concentration (ppm) trapped melt t f melt fraction (wt) t °c 0 0.2 0.4 0.6 0.8 0 200 400 600 800 1000 1200 1400 1600 1800 2000 2200 sh d* gabbros fe fe si fe si fe si fe si 0 0.5 1 1.5 2 fe si 1 1020106011001140 0 4 8 12 16 20 0 6 12 18 24 30 si lld: solid curves: si-enrichment dashed curves: fe-enrichment melt models: black curves: melt inversion gray curves: raleigh fractionation lzb uza uzb uzc st ra ti gr ap hi c he ig ht (m ) trapped melt t f melt fraction (wt) t °c 0 0.2 0.4 0.6 0.8 0 200 400 600 800 1000 1200 1400 1600 1800 2000 2200 sh 1 1020106011001140 0 5 10 15 20 25 0 5 10 15 20 25 30 35 40 0 10 20 30 40 50 60 70 80 0 1 2 3 4 5 6 7 8 0 160 320 480 640 800 0 500 1000 1500 2000 2500 si fe v 0 0.1 0.2 0.3 0.4 0.5 d* fe si 0 40 80 120 160 200 0 4 8 12 16 20 rb gabbros fesi raleigh fractionation si: 7.000 ppm fe: 10 000 ppm 0 0.2 0.4 0.6 0.8 1 0 400 800 1200 1600 2000 0 80 160 240 320 400 ba gabbros fe fe si si 0 0.1 0.2 0.3 0.4 0.5 0 2 4 6 8 10 0 0.2 0.4 0.6 0.8 1 u raleigh fractionation si: 265 ppm fe: 330 ppm si lld: solid curves: si-enrichment dashedcurves: fe-enrichment melt models: black curves: melt inversion gray curves: raleigh fractionation raleigh fractionation si: 14 000 ppm fe: 16 000 ppm fe mz lzc lza fig. 60 selected trace elements (sc, v, rb, ba, yb and u) from gabbros and modelled melt as a function of stratigraphic height (m) in the ls. first column shows the trapped melt and fraction of melt remaining (f) from tegner et al. (2009) and the liquidus temperature (t) calculated from thy et al. (2009a, 2013). for each selected trace element is shown the bulk distribution coefficients (d*, eq. (8)), trace element concentrations (ppm) in gabbros (data in table 13), inverted melt composition (eq. (9)), modelled melt composition assuming perfect rayleigh fractionation (eq. (11)) and initial melt compositions for the bottom of the column. the calculation of the individual melt/mineral distribution coefficients are described in section 5.11.1. two different solutions are shown for d* and all melt compositions, based on the two contrasting liquid lines of descent: prolonged iron enrichment past feti oxide appearance (fe) and iron depletion and silica enrichment following the appearance of feti oxides (si). the maximum concentrations in ppm for rayleigh fractionation at the sh are given for truncated diagrams (rb, ba, u). https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 99 of 139 geusbulletin.org high value (c. 45%) at the base of lza to c. 3% at the top of lzc (fig. 61; table 24). above the lz, trapped melt was shown to steadily approach a few percent in uzb. the uzc was not included in the study by tegner et al. (2009). the calculated di * coefficients in this study allow a re-evaluation of ftm including elements that are not highly excluded from the crystallising assemblage (fig. 61). the resulting variation in ftm with the stratigraphic height as averages for rb, ba and u resembles that of tegner et al. (2009) with an initial total amount of 32% for the lza, a systematic decrease upwards to 9% at the base of lzc, followed by a slight increase to 4% into the base of the mz. this is followed by a systematic decrease to 2–3% at the top of uzb, followed by a systematic increase in the sh (table 24a). in contrast to the low modelled ftm in the uzc, high residual melt components are suggested by the granophyre component widely interpreted as end-stage differentiation products for which crystallisation approaches ‘eutectic’ control (wager & brown 1967, p. 240). the modal estimates (fig. 15; table 2) point towards an increasing amount of modal quartz and orthoclase in uzc, reaching a total modal sum of 14% for quartz and orthoclase in the sh (table 2) that resides in a granophyre component of the gabbros (figs 13, 14). the modal composition of the granophyre patches and veins is mainly quartz, orthoclase and albite (wager & brown 1967; naslund 1989; hirschmann 1992). although quartz and orthoclase can be equated with the values in table 2, an amount of albite is included in the calculated plagioclase mode. wager & brown (1967, p. 138) estimated 10% of albite phenocrysts in a granophyre vein associated with the tinden sill. a better modal estimate of groundmass granophyre, based on normative composition, is perhaps provided by barker (1970) as 37% quartz, 29% orthoclase and 34% albite, approximating the granitic minimum (tuttle & bowen 1958). using this modal composition for the interstitial skaergaard granophyre and accepting the interpretation that the uzc granophyre represents an interstitial melt, it is possible to recast the modes for the samples collected as part of this study to include a realistic granophyre component reaching an average of 20% in the upper part of uzc (fig. 61; table 24b). the critical issue, however, is whether the granophyre component of the sh ferrodiorites can be taken as the end-stage melt for terminal fractional crystallisation approaching ‘eutectic’ control, as suggested by wager & brown (1967) and widely assumed to be the case. next, we consider the extent to which crystal fractionation processes can quantitatively account for the geochemical systematics observed in the ls and the possible effects of liquid immiscibility in the uz. 5.11.5 perfect fractional crystallisation the evolution of the trace elements is shown for selected included and excluded elements in fig. 60 for the two liquid descent models. the melt compositions were calculated in two different ways. firstly, the melt concentration was inverted from the gabbro composition using the bulk partition coefficient and eq. (9). secondly, the trace element concentrations were modelled assuming perfect rayleigh fractionation using eq. (11) and making some reasonable assumptions for the initial melt compositions at f = 1 or at c. 365 m depth in the hz. the similarities or differences between the two sets of calculated melt compositions can be used to evaluate whether perfect fractional crystallisation was a major operator during solidification. the variation seen in the inverted melt composition for included elements conforms reasonably well to the prediction from perfect fractional crystallisation (like sc and v in fig. 60). the lz often displays marked fluctuations around the predicted trend for perfect fractional crystallisation that can be attributed to uncertainties resulting from sampling the strongly heterogeneous layered gabbros (e.g. v). with advanced fractionation, the variation in st ra ti gr ap hi c he ig ht (m ) trapped melt (%) tegner et al. 2009 (p, rb, u) uzc uzb uza mz lzc lzb lza si-enrichment fe-enrichment + granophyre ftm averages of rb, ba, u 10 504030200 0 200 600 800 1000 1200 1400 1600 1800 2000 2200 400 fig. 61 modelled trapped melt (%) as a function of stratigraphic height (m) in the ls. model uses d* for selected excluded elements (average of rb, ba, u) and eq. (12), (table 24). two models are shown for si-enriched and fe-enriched melt, respectively. the results of the present study based on p, rb and u are compared to that of tegner et al. (2009). the effect of allocating the granophyre component in the uzc to the trapped melt is shown by a solid orange curve (labelled + granophyre; data from table 24). blue horizontal dashed lines: divisions of the ls. abbreviations for ls divisions in fig. 2. ftm: fraction of trapped melt. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 100 of 139 geusbulletin.org table 24 summary of preferred modes used for the trace element modelling including trapped melt zone stratigraphic height (m) melt ol opx cpx pl mt il ap kfs q high low (wt%) (wt%) (wt%) (wt%) (wt%) (wt%) (wt%) (wt%) (wt%) (wt%) a: calculated assuming perfect fractional crystallisation                   uzc 2165 2164 0.15 17.14 22.93 40.73 3.69 1.36 9.09 4.89 uzc 2165 2165 0.14 19.56 27.81 37.72 3.53 0.87 7.05 3.29 uzc 2163 2147 0.17 21.09 32.63 28.87 5.09 1.60 7.08 3.46 uzc 2144 2141 0.43 19.58 32.80 31.41 4.75 1.46 6.81 2.76 uzc 2091 2081 0.50 22.58 36.44 26.13 4.81 1.79 5.47 2.26 uzc 2075 2060 0.67 15.76 42.34 33.17 5.10 1.79 0.83 0.33 uzc 2060 2046 0.84 20.53 39.73 30.18 5.29 1.85 0.99 0.56 uzb 2044 2030 1.24 20.84 34.35 37.21 4.44 1.22 0.43 0.26 uzb 2024 1984 1.09 19.75 24.11 38.54 3.60 1.09 11.23 0.60 uzb 1962 1921 1.53 21.99 29.87 39.45 0.43 4.76 1.97 uzb 1902 1862 2.13 23.50 23.90 42.06 0.16 5.38 2.87 uzb 1843 1801 1.97 22.23 18.80 46.74 0.95 5.88 3.43 uzb 1782 1748 2.54 23.53 19.45 39.83 3.92 6.56 4.17 uzb 1728 1691 1.89 25.71 19.13 40.26 0.65 7.62 4.74 uzb 1671 1630 2.34 26.70 13.04 37.76 3.58 9.15 7.43 uzb 1620 1600 2.42 20.18 21.31 39.56 2.89 7.09 6.54 uza 1561 1527 2.68 10.56 32.68 38.97 5.39 6.75 2.96 uza 1506 1431 4.39 7.73 4.73 25.76 46.65 3.03 7.58 0.13 uza 1403 1343 2.85 8.85 0.81 31.12 44.64 2.49 9.17 0.06 uza 1323 1282 2.78 6.53 3.62 30.06 44.92 3.69 8.23 0.16 uza 1268 1237 2.89 11.48 0.00 31.65 40.59 4.98 7.92 0.48 uza 1218 1188 4.91 7.06 0.06 34.96 39.44 6.51 6.93 0.13 uza 1178 1131 4.32 3.79 1.24 37.95 40.21 3.06 9.24 0.19 mz 1098 1055 3.75 1.51 3.83 36.17 39.67 6.36 8.57 0.13 mz 1053 1032 2.23 0.22 8.12 24.82 45.78 4.39 14.33 0.10 mz 1024 1009 3.37 0.00 8.49 37.80 31.46 0.86 18.02 0.00 mz 1003 1002 3.03 3.37 7.80 34.49 36.12 5.37 9.76 0.06 mz 988 961 2.95 0.84 4.80 29.71 41.47 7.88 12.28 0.06 mz 953 930 3.33 4.03 3.80 41.19 26.18 3.87 17.51 0.10 mz 921 900 2.40 1.98 2.82 31.21 36.15 10.54 14.85 0.06 mz 890 875 2.62 1.69 5.97 27.79 47.99 4.37 9.50 0.07 lzc 847 816 2.83 3.75 0.03 29.73 41.39 6.88 15.33 0.06 lzc 808 798 2.41 6.00 2.90 27.18 29.45 10.11 21.44 0.51 lzc 784 742 4.78 5.19 2.61 22.61 30.23 6.55 27.84 0.21 lzc 723 703 8.02 1.91 6.84 39.66 31.45 4.13 7.90 0.09 lzb 703 681 6.23 6.20 8.19 32.12 43.60 1.67 1.88 0.12 lzb 634 580 9.99 4.38 10.75 24.71 45.85 1.36 2.34 0.63 lzb 558 488 10.74 3.11 7.87 29.77 46.22 0.35 1.57 0.36 lzb 447 367 13.37 7.90 10.16 23.98 42.54 0.22 1.69 0.14 lzb 346 221 12.99 5.69 13.11 20.34 45.71 0.40 1.65 0.11 lzb 177 173 21.28 2.17 11.59 19.81 39.67 3.58 1.82 0.09 lza 161 137 24.56 2.72 12.61 12.68 45.26 0.29 1.74 0.15 lza 125 107 27.73 5.74 11.99 10.53 41.46 0.87 1.46 0.22 lza 96 27 25.92 8.03 16.77 8.03 38.07 1.03 1.90 0.25 lza 7 31.06 9.34 11.71 9.42 36.33 0.46 1.47 0.21 b: uzc including adjusted granophyre content calculated from quartz, orthoclase, and part of plagioclase uzc 2165 2164 21.16 19.56 27.82 27.06 3.53 0.87 uzc 2165 2165 21.55 21.10 32.65 18.02 5.09 1.60 uzc 2163 2147 19.62 19.64 32.90 21.61 4.76 1.46 uzc 2144 2141 16.29 22.60 36.48 18.02 4.81 1.79 uzc 2091 2081 3.18 15.78 42.40 31.75 5.11 1.79 uzc 2075 2060 4.25 20.56 39.78 28.25 5.30 1.85 uzc 2060 2046 2.80 20.91 34.46 36.15 4.46 1.22 compare to table 2 showing the mineral modal variation without trapped melt. stratigraphic height is given as high and low referring to the interval used for averaging. ol: olivine. opx: orthopyroxene. cpx: clinopyroxene. pl: plagioclase. mt: magnitite. ap: apetite. q: quartz. kfs: orthoclase. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 101 of 139 geusbulletin.org inverted melt converges to the strongly depleted trend of fractional crystallisation that characterises ideal included elements reaching low levels of depletion late in the stratigraphic sequence (ni, cr and v). few of the excluded elements are truly excluded throughout the complete stratigraphic sequence. the elements shown in fig. 60 (rb, ba and u) are initially excluded, although with bulk partition coefficients often elevated by the inclusion of trapped melts. with advanced stages of fractionation, and the appearance of feti oxides, the bulk partition coefficients often exceed unity (zr and nb), particularly for the fe-enriched model for the liquid line of descent. the excluded elements, nevertheless, immediately show strong deviation from the perfect rayleigh fractionation curve, failing to reach sufficiently high melt concentrations. the only exception is ba, which shows reasonable correspondence between the inverted and fractional crystallisation melt compositions still into the uzb (fig. 60). the fractionation trends for three selected trace elements (rb, sc and yb) are illustrated in fig. 62 and compared to perfect fractional crystallisation for bulk distribution coefficients (di *). as previously noted, the included elements, like sc, reasonably conform to the rayleigh fractionation trend for variable values of di * (i.e. >1). this suggests only a modest underestimation in the di * values of fig 62 (1.2–2.0). on the other hand, the excluded elements, like rb and yb, strongly deviate from the rayleigh fractionation trends despite the predicted di * values for rb of about 0.1 and yb of about 0.4; for the latter, the saturation of apatite is followed by an increase to about 1.0 in uzc. although these di * values appear reasonable, the estimated melt compositions that use them are largely constant throughout the stratigraphic sequence, suggesting included di * values of 0.8–1.0. none of the examined elements correspond to the inverted and the predicted fractional crystallisation melts. however, the overall impression is that the included elements conform more closely to a first order fractional crystallisation model than the excluded elements. this supports a strong decoupling between the excluded and the included (major and trace) elements. these observations show that a unifying crystallisation process cannot explain the two groups of trace elements and that losses of significant amounts of excluded elements from the crystal mush prior to final solidification of the gabbros must be involved. 5.11.6 imperfect fractional crystallisation the melt composition evolves for perfect fractional crystallisation by removing or isolating early-formed crystals and thereby changing the melt composition in the crystallisation zone and eventually the bulk magma 0.2 0.3 0.4 0.5 0.6 0.9 1.0 0.90 1.0 1.1 1.2 1.5 2.0 3.0 0.2 0.3 0.4 0.5 0.6 0.7 0.8 1.0 rb 0 200 400 600 800 1000 1200 1400 1600 1800 2000 2200 0 20 40 60 80 100 melt concentration (ppm) rayleigh fractionation melt concentration (ppm) rayleigh fractionation sc 0 200 400 600 800 1000 1200 1400 1600 1800 2000 2200 yb d* d* d* st ra ti gr ap hi c he ig ht (m ) 0 10 15 200 10 20 30 40 5 melt concentration (ppm) uzc uzb uza mz lzc lzb lza sh inverted melt si enrichment fe enrichment 0.8 0.7 fig. 62 inverted melt concentrations of rb, sc and yb (ppm) compared to variable rayleigh fractions using the modelled d*values and ranges of constant d* values. si enrichment and fe enrichment refer to the two contrasting liquid lines of descent: fe: prolonged iron enrichment past feti oxide appearance, and si: iron depletion and silica enrichment following the appearance of feti oxides. blue horizontal dashed lines: divisions of the ls. abbreviations for ls divisions in fig. 2. solid black curve: perfect crystallisation (d* = 0). grey dashed curves: constant d* values (0.2 to 1). https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 102 of 139 geusbulletin.org chamber, only if the residual melt is recycled into the main chamber. perfect fractional crystallisation is nevertheless unlikely in a cooling magma chamber like the skaergaard intrusion, where components of equilibrium or imperfect fractionation may have been involved. the evolution of trace elements during ideal equilibrium crystallisation can be approximated using fractional crystallisation, eq. (10), assuming constant concentrations for each segment throughout the column fixed by its initial value f = 1. consequently, the initial melt composition will be approximately identical to the final composition because no solid material is removed from the crystallising system. calculations of imperfect fractional crystallisation are illustrated in fig. 63 for a variable component of equilibrium crystallisation and for three selected elements: rb and yb represent the excluded elements with low bulk partition coefficients, and sc represents the included elements with moderate coefficients. for the excluded elements, the inverted melt compositions correspond reasonably to the predicted ideal equilibrium crystallisation (rb and yb). for some excluded elements, like yb, the inverted compositions predict an apparent equilibrium component > 100%, likely to be attributed to small errors in d*. in contrast, the included elements, like sc, suggest a dominant mode of fractional crystallisation with only a modest equilibrium component at c. 40% in the mz, uza and into uzb. the poor correspondence, as observed for perfect fractional crystallisation, between the included and excluded elements again points towards the loss of a component of excluded elements, as observed in modelling of perfect fractional crystallisation. the modelling for sc, however, suggests equilibrium components of 40% in the mz and uza and high values at the top of uzb. furthermore, the excluded elements rb and yb imply near complete equilibrium that could have controlled crystallisation (fig. 63). this contrasts with the observation that all silicate minerals show strong and systematic variation in the overall stratigraphic column, proposing a dominant fractional crystallisation (figs 25–28). 5.11.7 in situ boundary-layer crystallisation modelling of fractional crystallisation implicitly assumes that primocrysts form homogeneously in the melt, controlled by phase equilibria and cooling, and are immediately removed and isolated in a bottom mush layer by various gravitative or convective processes. the magma that occupies pore spaces in the mush is thought to return to the convecting magma principally by compaction. studies rb 0 200 400 600 800 1000 1200 1400 1600 1800 2000 2200 st ra ti gr ap hi c he ig ht (m ) 0 20 40 60 80 100 melt concentration (ppm) equilibrium crystallisation 20% 40% 60% 80% 0% 0 402010 30 melt concentration (ppm) sc rayleigh fractionation equilibrium crystallisation 60% 20% 80% 0% uza mz lzc lzb lza uzc sh inverted melt si enrichment 0 105 15 melt concentration (ppm) 0 200 400 600 800 1000 1200 1400 1600 1800 2000 2200 rayleigh fractionation 80% 60% 20% 0% 40% yb uzb fig. 63 modelled melt concentrations (ppm) of rb, sc and yb for imperfect fractional crystallisation as a function of stratigraphic height (m) in ls. perfect fractional crystallisation is shown for comparison. curves are shown for 100, 80, 60, 40, 20 and 0% equilibrium crystallisation calculated for the si-enrichment liquid line of descent. blue horizontal dashed lines: divisions of the ls. abbreviations for ls divisions in fig. 2. solid black curves: perfect equilibrium crystallisation (100%). solid grey curves: perfect reaction crystallisation (0%). dashed curves: imperfect crystallisation (80–20%). https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 103 of 139 geusbulletin.org have nevertheless emphasised that heat loss and crystallisation of a cooling magma chamber may occur heterogeneously at the margins in narrow transition zones between the liquid and solid strata (e.g. jackson 1961; morse 1986; marsh 1988; tait 1988; jaupart & tait 1995; nielsen 2016). trace element concentrations of in situ boundary-layer crystallisation have been modelled by langmuir (1989) and nielsen & delong (1992). the principal difference between boundary-layer and homogeneous-chamber crystallisation is that for the former the melt in the pore spaces is returned to the main magma body, principally by compositional density convection, while in the latter, melt is compacted out of the mush. these differences in mush formation may not be distinguished petrographically, but the difference in the overall liquid line of descent may offer some clues (langmuir 1989). boundary-layer crystallisation implies that sub-liquidus stable phases may affect the overall fractionation as opposed to what would be expected from equilibrium liquidus relations. also, excluded element concentrations and ratios will be enriched more than included elements for in situ crystallisation compared to fractional crystallisation, and thus enhance the decoupling between excluded and included elements. the formulation for in situ crystallisation, eq. (13), proposed by langmuir (1989) is used in fig. 64 to illustrate the predicted variations for selected trace elements (rb, sc, and yb). curves for solid fractions (s) of 0.00, 0.25, 0.50, 0.75 and 1.00 are shown. the solid fraction of 0.50 is assumed to be a reasonable average for the skaergaard in situ crystallisation zone, as suggested by langmuir (1989) for the kiglapait intrusion. the ftm is calculated as shown in table 24, and the remaining melt is returned to the main chamber (for s = 0.50, the returned amount varies as a function of ftm between c. 0.25 and 0.50).the inverted melt composition for the two excluded elements rb and yb correspond to curves for constant and very low return of melt to the main magma (s << 0.25), while the moderately included sc corresponds to curves for high return in the lower and middle parts of the column (s = 0.75–1.00) to a decreasing 0 200 400 600 800 1000 1200 1400 1600 1800 2000 2200 0 20 40 60 80 100 melt concentration (ppm) rayleigh fractionation st ra ti gr ap hi c he ig ht (m ) rb 0 10 20 30 40 melt concentration (ppm) 0 5 10 15 melt concentration (ppm) sc yb rayleigh fractionation inverted melt si enrichment s sh s uzb uza uzc mz lzc lzb lza 0.750.50 0.50 0. 50 0.25 0. 00 1.00 0.00 0. 00 0.25 1.00 0.75 0. 75 1. 00 0. 25 s fig. 64 modelled melt concentrations (ppm) for in situ crystallisation of rb, sc, and yb as a function of stratigraphic height (m) in the ls. perfect fractional crystallisation and inverted melt compositions shown for comparison. trapped melt content and d* are based on the modes in table 24. grey dashed curves: solid fractions in the crystallisation front (s) of 0.25, 0.50 and 0.75, where sum of the solid fraction plus trapped melt and returned melt to the main magma is equal to 1. black solid curve for s = 1.0 is equivalent to perfect fractional crystallisation, while the grey dashed line for s = 0.0 displays constant variation equivalent to the initial concentration (co, see langmuir (1989)). blue horizontal dashed lines: divisions of the ls. abbreviations for ls divisions in fig. 2. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 104 of 139 geusbulletin.org return throughout most of the uz (s = c. 0.50). these diverse predictions for the fraction of melt returned for excluded and included elements confirm their strong decoupling as encountered for other crystallisation models and are not easily explained by crystallisation models (fractional or in situ). a low return fraction of melt (<0.25) to the main chamber will reduce or for an extreme condition (s = 0.00) prohibit in situ fractionation, while a high return fraction of melt will promote in situ fractionation effects. in the extreme, in situ fractionation for a complete melt return (s = 1.00) will be equal to perfect rayleigh fractionation (fig. 64). because of the low trapped-melt content modelled for the main part of the skaergaard ls, in situ crystallisation does not offer a solution to the observed variations in trace element. see sections 5.12 and 5.13 for a discussion of models for trapped melts accounting for effects of liquid immiscibility. 5.12 role of liquid immiscibility 5.12.1 evidence for liquid immiscibility the idea that liquid immiscibility affected the late-stage liquid evolution of the skaergaard intrusion is not new. it was first proposed by mcbirney & nakamura (1974) and mcbirney (1975) based on melting experiments using mixtures of granophyre and uzb and uzc ferrodiorites. this proposal was strengthened by the analysis and recognition of melts as inclusions in apatite, olivine and plagioclase (jakobsen et al. 2005, 2011), pointing towards an emulsion of si-rich melt and fe-rich droplets in the mushes and in the late stages of the magma chamber. experimental atmospheric support was provided by charlier & grove (2012) and honour et al. (2019a) on samples resembling potential skaergaard melts. further support was offered by veksler et al. (2007, 2008) by centrifuge experiments showing liquid immiscibility in skaergaard-like mz melts (see also veksler 2009; veksler et al. 2009). petrographic studies of interstitial material in the lower part of the intrusion have also invoked the actions of immiscibility melts to explain textural and chemical observations (humphreys 2009, 2011; holness et al. 2011). further, nielsen et al. (2015) have proposed a mechanism for the formation of the skaergaard pge-au mineralisation that incorporates migration of both ironrich and silica-rich immiscible melts in the upper part of the intrusion. there is thus little doubt that the formation of conjugate immiscible melts can and indeed did affect the skaergaard liquid line of descent and thus deserve careful consideration. following jakobsen et al.’s (2005, 2011) discovery of trapped emulsion droplets of melt in the uz gabbros, it was also expected that macroscopic and geochemical evidence for the action of liquid immiscibility during the solidification of the skaergaard intrusion would be found as part of this study. thus, the appearance of a high olivine and apatite base to the uzb and a high granophyre concentration at the top of uzc are seen as evidence for the action of immiscibility and the formation of an apparent zoned upper part of the uz (figs 15, 52c; see also sections 4.2 and 5.7). additional evidence for immiscibility is seen in the decoupling of included and excluded trace elements throughout the ls that we relate to the migration of latestage granophyre in the gabbro mush (see section 5.11). these observations suggest that liquid immiscibility may have affected the evolution during crystallisation of both the residual magma chamber and solidification of the melt trapped in the mush across the entire intrusion. this interpretation appears to contrast with the suggestion by nielsen et al. (2015) that immiscibility occurred in the mush and only indirectly affected the main chamber by buoyancy-controlled upward migrating granophyre. in this section, we develop a simplified intrusion-wide model for the effect of immiscibility on the liquid line of descent of trace elements that can explain the decoupling of included and excluded elements. we base this on the observations as well as reasonable assumptions of the behaviour of the trapped melt in the mush as well as the residual magma chamber. 5.12.2 liquid immiscibility and crystallisation in the residual chamber a summary of the model used here is shown in fig. 65 as simplified and schematic interpretations of the evolution of the skaergaard terminal magma chamber. it is suggested that magma convection in the residual chamber (fig. 65a) is terminated by liquid immiscibility (fig.  65b) and that partial buoyancy separation of the two cognate melts effectively results in a zoned chamber (fig. 65c) that with cooling crystallises as a mixture of ferrodiorites and granophyres (fig. 65d). the onset of immiscibility in the main residual chamber of the skaergaard intrusion is observed to have occurred at the uza–uzb boundary (fig. 50) at a temperature of c. 1060°c for a magma saturated in plagioclase (an40), olivine (fo39), clinopyroxene (mg# 50) and feti oxides (see section 5.7.7). we estimate that the sio2 and feo* contents at the onset of liquid immiscibility were 65% and 14% for the si-enrichment trend, while they were 54% and 26% for the fe-enrichment trend (tables 21, 22). we further suggest that convection in the main magma body may have ceased when large-scale liquid immiscibility developed, and the conjugate liquids separated gravitationally. the upward migration and coalescence of a granophyric component resulted in the host melt made of an apparent zoned coarse-textured https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 105 of 139 geusbulletin.org mixtures of ferrobasaltic (or dioritic) and silicic melts of which the latter eventually may have formed larger bodies as dykes or sills, transitioning into the lower part of the ubs or the host rocks (e.g. jakobsen et al. 2011; salmonsen & tegner 2013). the residual chamber would at that stage have been relatively small with an effective zoning impeding convective overturn. this would also have restricted or terminated large-scale fractional crystallisation and have resulted in an overall imperfect mode of fractional crystallisation mostly controlled by equilibrium crystallisation (see section 5.11). the two conjugate melts, during active immiscibility, would have been stable at an identical temperature (initially at c. 1060°c) and saturated in identical mineral compositions. following termination of immiscibility, because of cooling and migration of granophyre, the upward growth of the crystallisation front would continue, but now in more mafic melts, resulting in the formation of the olivine-rich and apatite-rich base of the uzb (fig. 15) and the trapping upward of increasing higher amounts of granophyre in a ferrobasalt (or dioritic) matrix (cf. fig. 14). the two components, although having cooled down simultaneously, would, however, have behaved differently due to markedly different liquidus and solidus temperatures and phase equilibria. compositional variations throughout uzb and uzc would then have been a result of the relative proportions, crystallisation, and reactions between the various melt components and not a consequence of protracted crystal fractionation. the details of these processes are based on our current knowledge and understanding of the effects of liquid immiscibility on the skaergaard magma but are not predictable and will require future experimental and detailed petrographic studies. ls/uza sh ubs/uza’ ls/uza sh ubs/uza’ uzb uzb’ b 10 60 °c ls/uza sh ubs/uza’ uzb uzb’ c fe-rich melt si-rich melt 10 60 °c convective magma uzb’ uzb liquid immiscibility a >1 06 0° c uzc uzc’ ls/uza sh ubs/uza’ uzb uzb’ d ~1 00 0° c ferrogabbro granophyre fig. 65 schematic illustration of the proposed late history of the terminal skaergaard magma chamber (uzb and uzc). a: uza convecting chamber at just above 1060°c and a melt fraction of 0.10, sandwiched between the upward-moving uza of the ls and the downward-moving uza’ of the ubs. b: onset of liquid immiscibility when the temperature reaches 1060°c by the formation of basaltic iron-rich droplets in a homogeneous silicic melt, effectively terminating chamber convection. c: separation of the two cognate melts with the si-rich melt preferentially migrating upward and the fe-rich melt droplets migrating downward (or remaining stationery). this results in an effectively zoned residual chamber with the fe-rich melt component concentrated principally at the base and the si-rich melt concentrated at the top under the uza’. immiscibility is subsequently terminated by the melt separation and the drop in temperature to unknown values. d: final crystallisation products at solidus temperatures (c. 1000°c and f = 0) are made up of a ferrodiorite (andesine, olivine and pyroxene) with granophyre (albite, orthoclase, and quartz) dispersed dominantly in the uzc and uzc’ from a microscopic interstitial to macroscopic lenses and sill scales. the granophyre component dominates the ubs. part of the ubs is thought to have been intruded into the surrounding host rocks. illustrations are not to scale. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 106 of 139 geusbulletin.org 5.12.3 crystallisation and immiscibility in the trapped melt the foundation for our model calculations is that the interstitial trapped melts crystallise in a similar fashion as seen for the overall intrusion. this is a simplified view of the mush pile because it considers crystallisation in the mush to occur by fractional crystallisation in a closed system without, or with limited, melt migration laterally or horizontally in the mush. however, this simplification allows estimates of the initial liquidus temperature based on plagioclase composition (thy et al. 2013), the onset of immiscibility assumed to be at 1060°c, and the amount of liquid remaining in the mush melt zone of about 0.10 of the initial trapped melt (tegner et al. 2009). it is possible that the trapped melt solidified with a significant component of equilibrium (imperfect fractional crystallisation), reactions with the host gabbro, or local redox conditions that allowed the trapped melt to either delay or accelerate the point of immiscibility compared to the overall differentiation processes inferred from the main chamber. thus, detailed petrographic studies of mush gabbros (humphreys 2009, 2011; holness et al. 2011; nielsen et al. 2015) have suggested that the trapped melt may have reached immiscibility prior to the similar event in the main magma chamber without, or with limited, exchange of granophyre from the mush to the main chamber. this is, however, not suggested by this study that observes macroscopic evidence for migration of both the basic and silicic cognate melt components in the uz chamber at temperatures similar to those likely experienced by the mush (see section 5.7.7). the modelling of the mush and chamber melt based on identical crystallisation conditions are nevertheless a simplification adopted to predict onset of immiscibility and its effects on the liquid evolution. if the trapped melt crystallised due to perfect fractional crystallisation without reaching immiscibility to a final melt also of a granophyric or granitic composition, the liquidus temperature would have been c. 950°c or below. the amount of this residual melt would have been restricted, and it would have been difficult for the melt to buoyantly segregate from the solidification zone. encountering the two-liquid field earlier during crystallisation of the mush liquid would have provided greater opportunity for buoyant segregation of the conjugate felsic liquid from the solidification zone at the floor of the intrusion. the porosity and permeability at the point of immiscibility is dependent on the fraction of trapped melt or the mush porosity and the physical properties of the melts. the high fractions of trapped melt mush in lza and lzb, reaching 1060°c, would constrain the porosity to a fraction of 0.01–0.04, which may have allowed interconnected permeable flow (turcotte 1982; mckenzie 1984, 1989). however, the greatly reduced melt fractions in the mush of lzc to uza (fig. 61), as well as locally, may still have acted as a barrier to granophyre migration, thus effectively trapping it in situ or only allowing lateral migration and coalescence into lenses, perhaps tectonically induced in the mush packet. the solidification of the lz mush mostly occurred prior to the onset of liquid immiscibility in the residual interstitial melt (tegner et al. 2009), allowing for a crystallisation interval from c. 70°c at the base of the exposed lza and linearly decreasing upward to the top of uza, where a liquidus temperature of 1060°c (f ≈ 0.10) is reached in the main chamber, and immiscibility is inferred to be initiated. in the lz and mz residual mush, immiscibility is therefore reached well after the main crystalline framework of silicate and oxide minerals has been established and the immiscible products thus completely or partially trapped. late-stage migration of granophyric melts may therefore mainly be controlled by ductile shear stress and shrinking fractures in the solid mush and to a lesser extent by compaction (larsen & brooks 1994; jakobsen et al. 2011; salmonsen & tegner 2013). this results in pods and veins of granophyre and melanogranophyre in the gabbros. approaching the uza–uzb boundary, where the solidification interval would trend towards zero thus allowing for more interactions between crystallisation and immiscibility and thus for a higher proportion of upward migration of granophyre melt (modelled from 0% in lza to 99% at the uza-uzb boundary). continued growth of the crystallisation front, following initiation of unmixing, upward and inward, is thus likely to trap conjugate melts or emulsion in larger amounts than lower in the stratigraphy (jakobsen et al. 2011). the silicic granophyric melt is less dense than the basic conjugate melt and the crystal network (fig. 59; hunter & sparks 1987; tegner et al. 2009). it is consequently probable that the silicic immiscible melt is able to partially migrate upward (nielsen et al. 2015) and that its success in reaching the main magma chamber will increase from minor to unlikely in the lz to probably successful in the uza. the accumulative fraction of the silicic conjugate melt, assumed to be drained from below into the uzb magma reservoir, was always small because of the low content of trapped liquid and the likely interruption of migration paths. small fractions of granophyre will have detectable effects on the uz evolution only for elements such as rb that are strongly partitioned into the silicic immiscible melt but local effects may still be detectable on the interstitial mineralogy and mineral compositions, as demonstrated by holness et al. (2011) and humphreys (2009, 2011). the role of the primitive to evolved basaltic melt filling interstitial spaces in the mush is an unknown factor and unaccounted for in the present modelling. the https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 107 of 139 geusbulletin.org physical properties and the mush porosity will control such melt migration prior to immiscibility. because of the solidification temperature interval and upward decrease in residual melt volume, the effects of basaltic melt migration will be most pronounced in the lower part of the stratigraphy. the density of the interstitial basaltic melt is slightly lower than the coexisting primocryst plagioclase until the base of the uza where the buoyancy reverses (fig. 59). the melt density is also much lower than the bulk-mineral network throughout the stratigraphy. thus, the basaltic melt could be prone to migrate upwards and laterally under appropriate porosity conditions and particularly if compaction or lateral shear stress were operative (tegner et al. 2009). gabbroic segregation lenses, formerly referred to as “wavy pyroxene rocks” by wager & deer (1939), occur in the mbs where they have been attributed to tearing and sagging in the vertical crustal mush (humphreys & holness 2010). larsen & brooks (1994) did an extensive examination of gabbroic pegmatites with geometrical shapes from discordant pods and domes in the lz and mz, transitioning into concordant lenses in the uz. they attributed these to segregated and buoyantly-migrated basaltic melt that crystallised under final hydrous conditions to mimic the overall fractionation in the intrusion, including a terminal immiscible granophyre melt. it is unfortunately not possible to quantify the volume of basaltic melt that may have migrated in the mush. the effect would be to reduce the absolute volume of immiscible melts and a reduction in excluded trace elements in the final gabbros. this latter effect is considered minor and would not have strongly affected the qualitative findings due to the much higher concentration of excluded elements in the immiscible granophyre melt. 5.12.4 predicting the effects of liquid immiscibility the observed decoupling of included and excluded elements suggested from the forward modelling of the trace elements was related to the effects of the migration of immiscible granophyre melts in the mush or an underrepresentation of granophyre in the analysed mush gabbros (see section 5.11). the immiscibility model formulated here allows us to evaluate the effects of this migration or loss of cognate melts. for this purpose, we use the observed gabbro column as described by composition (figs 33–35), modal and mineralogical makeup (fig. 15), predicted partitioning coefficients and estimated trapped-melt content (fig. 61). as noted, we assume that immiscibility reached the liquidus at 1060°c and that 10% of the initial trapped melt remained, as for the bulk chamber melt. it should, however, be considered that the model calculations only consider the endstage evolution of the gabbro column and its cooling to the point when liquid immiscibility occurs. thus, it does not consider the mode of formation of the final mush, either by in situ crystallisation, compaction or a combination of these two. the purpose of the model calculations is to (1) quantify the behaviour of the uzb and uzc, and (2) quantify the possible effects of the horizontal and lateral migrations of the silicic immiscible melt in the lza–uza gabbro column. the trapped-melt fraction in the gabbro mush is known, or believed to be known, from the present estimates (table 24; fig. 61). this means that at the point miscibility reaches the liquidus, the conditions and compositions can be taken throughout the lower gabbro column to be identical to those at the uza–uzb boundary. we can thus predict at individual levels in the ls, as well as cumulatively, the amounts by weight of granophyre liquid that potentially may have been affected by migration, the trace element content prior to immiscibility, the trace element contents of the two conjugate melts, and temperature (taken to be a constant at 1060°c). the basic assumptions are obviously that the trapped-melt content is correctly estimated and that the sampled gabbros represent the final crystalline mush (see section 5.11.4). the trace element concentrations at the point of immiscibility in both the mush and the melt chamber can be estimated by rayleigh fractionation modelling (fig. 62). further, the trace element compositions of the two conjugate melts can be predicted based on experimental partition coefficients between the immiscible silicic and basaltic conjugate melts for which the partitions are reasonably known (rb, sc and ba for which data are available; watson 1976; watson & green 1981; veksler et al. 2006; veksler & charlier 2015). rubidium is dominatingly partitioned into the silicic melt (ds/b = 5), while ba is about evenly partitioned (ds/b = 0.8), and sc is moderately partitioned into the basic melt (ds/b = 0.5), where ds/b is partition coefficient between immiscible silicic (s) and basic (b) conjugate melts. the partition table 25 variables used for modelling of liquid immiscibility elements rb ba sc partition coefficients d (silicic (s) / basaltic (b) conjugate melts) ds/b 5.0 0.8 0.5 db/s 0.2 1.3 2.0 bulk partition coefficients d* (solid/liquid) d* uza-uzb 0.16 0.15 7.9 d* mz-uza 0.16 0.18 3.7 d* lza 0.4 0.4 0.8 element concentrations (ppm) at uza–uzb boundary rayleigh fractionation 74 798 0.07 silicic conjugate melt 62 363 0.02 basic conjugate melt 12 436 0.04 t (ºc) 1061.5 f 0.105 https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 108 of 139 geusbulletin.org coefficients and the concentrations of the selected trace elements in the original melt and in the two conjugate melts are shown in table 25. the only other information needed is the relative proportions of the two conjugate melts. most experimental studies have argued that the proportions of the two melts might reach a weight ratio of one (dixon & rutherford 1979; veksler et al. 2007, 2008; charlier & grove 2012; charlier et al. 2013). veksler et al. (2006, 2008) examined a ferrobasaltic composition, similar to the major element composition used here, to determine the trace element partitions and obtained about equal proportions of the two conjugate melts. in contrast, our results (fig. 53) suggest a higher proportion of the silicic (c. 80–70%) compared to the basic (c. 20–30%) conjugate melt, depending on the liquid line of descent model considered. these estimates, however, are based on forward fractional crystallisation modelling in contrast to the equilibrium experimental results. a high amount of equilibrium conditions can be expected to result in a marked steepening of the liquid decent line. for this reason, we model below both the fractional crystallisation (80% silicic melt) and the equilibrium crystallisation (50% silicic melt) modes. it is further assumed, because of the lack of knowledge, that the proportions of the two immiscible melts remain constant with decreasing t < 1060°c, until the liquidus emerges from the immiscibility field at an unknown temperature. the basic immiscible melt both at levels below and above the uza–uzb boundary is, with minor exceptions, taken to be fully retained where it formed in the mush. it is only assumed that small amounts of basic conjugate melt in the uzb and uzc may partially sag in the main chamber and that likewise a small component of the ubs basic conjugate melt may also migrate down into the uz chamber (nielsen et al. 2015). although some of these assumptions might be questioned or refined, they will serve as first-order constraints, allowing us to predict the general effect of liquid immiscibility, although not as much the absolute effects. fig. 66 shows three possible 1060°c chamber models considered in the modelling of the trace elements as variations in the melt fractions of the two conjugate melts (basic and silicic) with stratigraphic height. the lower part of the intrusion is not shown because of the predominantly very low conjugate melt fractions. the first two models (in figs 66a, b) assume a significant amount of granophyre migration out of the chamber into the ubs and host lavas. fractions of lost granophyre are taken to be 0.38 (fig. 66a) and 0.56 (fig. 66b); volumetrically equivalent to 8.5 km3 and 12.4 km3 on a weight basis, respectively, using nielsen’s (2004) volume estimates. the models differ in the assumed ratio between the silicic and basic conjugate melts: the model in fig. 66b is scaled as 80% silicic melt and 20% basic melt, where as in fig. 66a it is 50% silicic melt and 50% basic melt. this scaling of silica to basic melt approximates to fractional crystallisation and equilibrium crystallisation, respectively (fig. 53). the model in fig. 66c is similar to that in fig. 66a, except that granophyre is not migrating out of the melt chamber. also included is a small 1600 2000 st ra ti gr ap hi c he ig ht (m ) 0 0.30.20.1 0.50.4 0.70.6 1800 2200 1600 2000 1800 2200 1600 2000 1800 2200 a b c immiscible melt fractions b s s s b b s b s s s s b b s silicic melt 50% b; 50% s loss of granophyre 20% b; 80% s loss of granophyre 50% b; 50% s no loss of granophyre b basic melt fig. 66 liquid immiscibility mass balance models for uzb and uzc. a: loss of granophyre from the intrusion (red) assuming 50% silicic and 50% basic conjugate melts. b: loss of granophyre from the intrusion (red) assuming 80% silicic and 20% basic conjugate melts. c: no loss of granophyre from the intrusion assuming 50% silicic and 50% basic conjugate melts. all three models involve a small migration of granophyre from the gabbro package below (grey dashed curve) and sinking of basic melt from the ubs (black dashed curve). for details of the mass balance models, see section 5.12.4. abbreviations are: b: residual basic conjugate melts. s: residual silicic conjugate melts. grey arrows: component migration down. red arrows: component migrating out of chamber. black arrows: component migrating up from below. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 109 of 139 geusbulletin.org amount of basic conjugate melt seeping down from the upper crystallisation front (i.e. the ubs) amounting to a fraction of 0.12–0.04 km3 or 2.8–0.9 km3, dependent on the ratio of the conjugate melt used in the model. a final part of the model is a small amount of granophyre seeping up from the ls into the chamber that amounts to 0.7 or 1.1 km3 and is similarly dependent on the ratio of the two conjugate melts. the result is a modestly zoned uz chamber, being either dominantly basic (models in fig. 66a and c) or silicic (model in fig. 66b) mainly dependent on the behaviour of the granophyre melt component. because little quantitative information is available on both the relative proportions of the two cognate melts as well as their early and late-stage behaviour in the mush and magma chamber, the purpose of the three upper-chamber models is principally to test the general effects of (1) loss of granophyre from the intrusion, and (2) the proportions of the two melts formed from immiscibility. there are nevertheless some first-order observations that can be made. to produce the observed mafic base of the uzb (figs 51, 52c), a high proportion of the basic conjugate melt is required – approaching 50% of the original melt. this allows us to exclude the model in fig. 66b that would produce a consistently silicic zoned terminal sequence of the skaergaard intrusion, and instead focus on the models of fig. 66 a and c. this in effect means that the modelling uses equilibrium crystallisation and not fractional crystallisation for the final evolution, as also predicted from the trace element variation. it is important to realise that although zoning is used to describe the final chamber of the skaergaard, this zoning is sensu lato caused by the increasing amount of a granophyre component towards the top and not necessarily by a sensu stricto zoned melt component. the amount of granophyre extracted in the models from the skaergaard magma is 8–9 km3 (model in fig. 66a), which amounts to 4–5% of the whole ls and 38% for only the combined uzb and uzc, based on the total volumes estimated by nielsen (2004). the question is obviously if this amount is a reasonable estimate for the ls contribution of granophyre to the ubs and the nearby host volcanics. the bulk composition of the intrusion estimated by mcbirney (1989a) did not incorporate granophyre or melanogranophyre. for this reason, nielsen (2004, table 3) added 5% of melanogranophyre to mcbirney’s (1989a) bulk composition for it to be normatively comparable to the tholeiitic lavas of the east greenland plateau (see section 5.5). this estimate implies that a volume of 15 km3 (melano-) granophyre could have potentially made up the final differentiate in the uzc, sh and the equivalent parts of the ubs. attempts by gabbro summation to estimate the bulk (or parental) magma compositions as well as the liquid line of descent have variably added granophyre or melanogranophyre amounting to 2% in lz–uza, 5–8% in uzb and 15% in the uzc (including their ubs and mbs equivalents; tegner 1997; nielsen et al. 2009; andersen 2006). in these calculations, melanogranophyre was mostly added instead of granophyre because these former compositions were judged more reliable. the accumulation of such additional volumes is proportionate to the lower part of the stratigraphy (5.5 km3), to the upper part of the stratigraphy (2.4 km3) and to the total intrusion (7.9 km3). the uzc equivalent part of the ubs (uz’) was mapped and described by naslund (1984) who observed that the lowermost part of the ubs was enriched in several low-melting and volatile components (sio2, k2o, and h2o), and hypothesised that these enrichments could be the result of upward migration from the underlying chamber and accumulation along the downward crystallisation front. salmonsen & tegner (2013) detailed the same transition in their kilen profile and estimated the thickness of the uzc’ to be 100 m or 30% of the total uz’ (or ubzγ3 and ubzγ, respectively, using the nomenclature of naslund 1984). compared to the corresponding ls, the uzc’ is relatively thicker in the central profile than the stratigraphy of the uzc, occupying 110 m or about 10% of the total uz. these observations may suggest a relatively thicker package of gabbros on the downward moving crystallisation front of the ubs and support the hypothesis that the uzc’ received late differentiation components like granophyre from the final static chamber. the final differentiation products (uzc–sh– uzc’) were estimated by nielsen (2004) to be distributed over an area of about 30 km2 in the inner parts of the intrusion or about 6 km3, evenly distributed between the two stratigraphic parts – these are somewhat high compared to the 4–5 km3 estimated by nielsen (2004) for the uzc alone. the petrographic and structural make-up of the final differentiate in the uppermost subzones around the sh was not well documented (wager & brown 1967). felsic differentiates of granophyre or melanogranophyre appears as pods and lenses in the uzc amounting to an estimated 15% of the gabbros (mcbirney 1989a). tegner (1997) noted that melanogranophyre lenses increased upward in the uz towards the sh, where they amounted to 15% of the total subzone. these estimates are similar to this study suggesting approximately 20% granophyre content towards the upper part of the uzc. the description of irvine et al. (1998) closely matches our impression that the uzc “[..] is a pegmatitic cumulate with plagioclase, olivine, hedenbergite, apatite, and magnetite being recrystallised cumulus minerals, and granophyre forming a postcumulus mesostasis.” in the uz’ of the ubs, the amount of granophyre is believed to be higher. although quantitative estimates of their spatial distribution are not available, an estimate of 25% https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 110 of 139 geusbulletin.org appears reasonable when studying the descriptions and chemical analyses of salmonsen & tegner (2013). these estimates of granophyre content of the uzc and uzc’ (a total of 1.2 km3) are well below the total amount predicted (6–9 km3), suggesting that large amounts of granophyre may have been lost during the final crystallisation stages of the skaergaard intrusion and may now reside in some of the associated granophyre lenses and sills. 5.12.5 observed, predicted and ideal liquid lines of descent the effects of immiscibility and selective migrating of the granophyric component in the mush pile of the lz to uza will markedly affect the observed liquid descent. in the following, we refer to three different calculated decent lines: ‘observed’, ‘predicted’ and ‘ideal’. the information obtained from the gabbro compositions are ftm and d*. these allow us to calculate the observed liquid descent line as the inverted melt compositions using eq. (9). the ideal decent without granophyre loss can be calculated by perfect fractionation using eq. (11). the lower part of the intrusion is modelled by perfect fractional crystallisation (lz–uza) and the upper part of the intrusion by equilibrium (or imperfect fractional) crystallisation (uzb–uzc; see section 5.12.2). the predicted, or restored, liquid lines of descent can then be calculated by the perfect fractional or equilibrium crystallisation processes using eq. (11) adjusted for the effects of formation of conjugate granophyre melt retained or lost from the mush as described below. the trace element concentrations in the conjugate melts are constants (table 25), while the amount of predicted fractionation without granophyre loss observed inverted melt ideal perfect fractionation with granophyre loss observed inverted melt observed inverted melt ideal perfect fractionation predicted fractionation id ea l p er fe ct fr ac tio na tio n observed inverted melt with granophyre loss without granophyre loss predicted fractionation 0 200 600400 ba (ppm) 0 302010 40 sc (ppm) 0 302010 40 uzb uza uzb uza 0 400 800 1200 1600 2000 st ra ti gr ap hi c he ig ht (m ) rb (ppm) fig. 67 immiscibility models of rb, sc and ba (ppm) shown as a function of stratigraphic height (m) in the ls. the observed inverted melt composition (solid grey curve) is for mass balance model a and c from fig. 66. the ideal perfect rayleigh fractionation (dashed grey curve) and the predicted fractionation (black curve with grey dots) are shown. the latter assumes that granophyre is lost from the ls column to either the upper zone chamber, or by migration out of the system to the ubs and into the intrusion host rocks or that the samples used here did not include granophyre. the lower part of the ls is modelled by rayleigh fractionation (lz–uza) and the upper part as equilibrium crystallisation (uzb and uzc). https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 111 of 139 geusbulletin.org trace elements lost from or added to the mush depends on the amount of trapped melt or the proportions of the conjugate melts. the predicted liquid lines of descent are thus an attempt to duplicate the observed compositions for which reasonable similarities or differences between these two decent lines reflect the success or failure of the calculations. the granophyre component in the lower segment of the intrusion (below uzb) is considered to (1) be retained in the matrix of the gabbro, (2) migrate upward and enter the chamber as shown in fig. 66, and (3) migrate laterally and coalesce in lenses and thus be lost from the mush by a non-mush representative sampling of the gabbro. we cannot distinguish between types (2) and (3) and thus group these together as lost or migrated granophyre. however, because we know the trapped-melt content (table 24), we can use the constants in table 25 to predict the ratio between the retained and lost granophyre and use a trial-and-error procedure to improve the model fit. in contrast to the ls below uzb, granophyre within the uzb and uzc magma-chamber granophyre migrated upward by variably increasing from zero at the base of the intrusion to 99% at the base of uzb. a portion is thus retained, and the rest is lost or otherwise not represented by the gabbros. fig. 67 shows the modelling results for three trace elements (rb, sc and ba), based on the arguments outlined above by incorporating the effects of both interstitial liquid immiscibility and crystallisation (fig. 65; table 25). the lower part of the intrusion is modelled by rayleigh fractionation (eq. (11)) and the upper part of the intrusion by equilibrium crystallisation. both segments were, in contrast to the traditional approach, modelled using a partially open system with a late-melt composition that is continuously changing due to liquid immiscibility-controlled migration of granophyre melt. rubidium is preferentially partitioned into the conjugate silicic melt (granophyre) during immiscibility and into the melt phase during fractional crystallisation (table 25). the results for the lower sequence (lz–uza) are thus a good fit between the observed melt and the predicted compositions, while the ideal rayleigh fractionation descent line deviates strongly from both the inverted melt and the predicted melt compositions (fig. 67). scandium behaves in the opposite fashion of rubidium as it is preferentially partitioned into the immiscible basic melt and into the solid fraction during crystallisation (table 25). because of the very low scandium content, migration of granophyre will have indiscernible effects. the result is a reasonable, although not perfect, fit between the three observed and predicted descent lines (fig. 67). the differences between the observed and the predicted fractional crystallisation may in the upper part of the sequence be attributed to the very low concentrations and high analytical uncertainty of scandium. barium is approximately evenly distributed between the two conjugate immiscible melts (table 25) and, like rubidium, is strongly partitioned into the melt phase during crystallisation. the descent lines for barium show a reasonable fit between the observed and predicted curves only if a higher fraction of granophyre is retained in the gabbros (60%) than for rubidium (20%). it is possible that these marked differences may be related to shortcomings in the partition coefficients for barium. because the uzb–uzc chamber is modelled for equilibrium crystallisation, the scenario without granophyre migration would be expected to show little upward variation in trace elements in contrast to a loss of granophyre that would show an upward decrease in the elements that are preferentially partitioned into granophyre. this is consistent with rubidium that corresponds to the modelling for granophyre loss because of a preferential partitioning into granophyre. barium evenly partitioned into the two cognate melts shows small differences between the models for loss of granophyre or without loss (fig. 66). the pooled granophyre loss from the ls amounts to 192 ppm rb and 1158 ppm ba (sc < 1 ppm). considering the sampling limitations (table 18), these concentrations are comparable to some skaergaard granophyre veins (larsen & brooks 1994; mcbirney 1989a; salmonsen & tegner 2013), transgressive granophyres (hirschmann 1992) and the basistoppen granophyre (naslund 1989). lastly, most studies have assumed that the two conjugate melts were unable to separate and that an emulsion persisted to temperatures well below 1060°c (veksler et al. 2008; jakobsen et al. 2011; charlier & grove 2012). the present study indicates partial separation and thus does not support a persistent emulsion in the uz, where liquid immiscibility may have contributed markedly to the differentiation processes. several trace elements show a strong exponential increase approaching the final crystallisation fraction of the sh (e.g. sc, rb, zn, ba, ree, u; figs 34, 35). some of these terminal increases appear to be consistent with fractional crystallisation (rb, ba), while others may not (sc; fig. 67). mcbirney (1995, 1996) argued that a latestage buoyant liquid that is rich in volatiles and excluded elements rose from the upper part of the intrusion and permeated above the sh into the ubs to form a shadow sh horizon located slightly above the sh. mcbirney and co-workers traced the origin of this to an observed decrease in cl in apatite at the base of uzb and proposed the exsolution of a hydrous and chlorine fluid (sonnenthal 1992; mcbirney 1995) at about the level where liquid immiscibility was initiated. mcbirney (2002) discussed the same observation and suggested that the https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 112 of 139 geusbulletin.org increase in excluded elements was caused by a gradual change in partition coefficients due to volatile-saturated melts from the top of the mz. however, tegner et al. (2009) argued that the action of a cl-rich fluid phase would have strongly fractionated rb from u in contrast to the observations, and thus rejected any involvement of a fluid-saturated melt in the uz. 5.12.6 effects of granophyre migration on calculated trapped melt content by incorporating the predicted granophyre loss, the restored gabbro compositions will markedly affect the calculated ftm. this restored trapped-melt content is shown in fig. 68 together with the corresponding bulk partitioning coefficients calculated by an iterative procedure and eq. (12) for rubidium concentrations of the granophyre adjusted gabbro compositions (drb *). the d* values for rubidium (and for sc and ba) were also used to calculate the liquid lines of decent shown in fig. 67. the first-order result is a marked rise in ftm upward, but in a pattern similar to that observed for calculations without granophyre adjustments (fig. 61; table 24). the adjusted trapped-melt content, reflecting the drb * variation, is about 40% in the lza, decreases to 6% in the uzc, and subsequently increases to 20% at the top of the uzb, followed by an increase to high values at the sh (fig. 68). we also observe that ideal rayleigh fractionation at the uza–uzb boundary reaches concentrations of about 25 ppm rb, compared to about 80 ppm without granophyre adjustment. the subsequent crystallisation in the uzb and uzc does not change the rubidium concentration, which at the sh reaches an end-point concentration of about 30 ppm. tegner et al. (2009) assumed that the fractionation throughout the ls was controlled by rayleigh fractionation and thus appear to have implied that the final product was granitic or granophyric. our study suggests that the uzb and uzc crystallised by a process more similar to equilibrium crystallisation (or imperfect fractional crystallisation; fig. 65). a similar observation was indirectly made by namur & humphreys (2018) based on primocryst-excluded trace element modelling. they observed good fits in their trace element modelling for uzb and uzc only if they used a trapped melt content ftm n o tr ap pe d m el t n o tr ap pe d m el t ideal perfect rayleigh fractionation uzb uza uzb uza predicted fractionation st ra ti gr ap hi c he ig ht (m ) observed inverted melt ideal perfect fractionated melt percent of granophyre added uzb uza 0 0.40.2 0.60 0.2 0.4 0.6 0 400 800 1200 1600 2000 5 10 15 350 20 25 30 a b c drb* rb (ppm & %) fig. 68 calculation of fraction of trapped melt (ftm) as a function of stratigraphic height (m) in the ls, using mobile granophyre components and an adjustment of the fractionation modes. a: observed inverted melt content of rb (ppm), ideal perfect fractionated melt content of rb (ppm), and the percent granophyre added to obtain the ideal fractionation curve. b: bulk partitioning coefficients for rb (drb *) compared to drb * values without trapped melt. c: the predicted ftm obtained by adjusting the gabbro mush composition, restoring the gabbro to a mush composition after adding granophyre. the granophyre used in the modelling contains 40 ppm rb (table 18). https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 113 of 139 geusbulletin.org of c. 60% (as opposed to c. 5% according to tegner et al. 2009). this would drive the bulk-partitioning coefficient towards unity and supports a strong component of equilibrium crystallisation. their observation of a high component of equilibrium crystallisation are not substantially different from the findings for the uz in the present study. this would mean that gabbros in the upper part of the intrusion (uzb and uzc) change little in composition and thus allowed the final consolidated rocks to become a mixture of ferrodiorites mixed with migrated and coalesced granophyre lenses resulting from a uzb liquid immiscibility and the possible reactive products from the two components. several studies have attempted to quantify the thickness of the mush layer at the solid-melt interface using petrographic and microstructural observations for the skaergaard intrusion (namur et al. 2014; nielsen 2016; holness et al. 2017b, d; vukmanovic et al. 2018) and the bushveld complex (holness et al. 2017c, d; kruger & latypov 2020). namur et al. (2014) showed that plagioclase-zoning profiles reflected the physical properties of the crystal mush and its porosity. holness et al. (2017b) observed no correlation between estimated trapped melt and crystal plastic deformation in the centre of the skaergaard intrusion, leading them to propose that compaction had been of minor importance and that the magmatic fabric reflected a thin layer of primary mush. holness et al. (2017b) further evaluated the mush thickness at the appearance of apatite at the uza–uzb boundary. they used microstructural stepchanges in the latent heat accompanying the appearance of apatite to suggest that the mush thickness in the centre and eastern parts was only a few metres thick, increasing to c. 100 m in the western part of the intrusion (all at the uza–uzb boundary). similar studies of the rustenburg layered suite of the bushveld complex (holness et al. 2017d) suggest a mush thickness of the order of a few metres and that neither gravitationally driven compaction nor compositional convection were likely to have been important processes during solidification. it was also suggested that adcumulates are most likely formed at the top of the mush during primary crystallisation. furthermore, gabbro blocks were believed to have detached from the roof zone and sink through the melt reservoir to rest on the contemporaneous mush (irvine et al. 1998; sonnenthal & mcbirney 1998). these blocks range from a few centimetres to several hundred metres in size. they are regarded to have indented the bottom mush on impact and are overlain by subsequent draping layers. this suggests a well-defined interface between the melt reservoir and the solid gabbro substratum and a thin mush zone (irvine et al. 1998). in conclusion and paraphrasing holness et al. (2017d) and holness (2018), the main issue with some of these arguments is that they are often based on bulk geochemistry and an arbitrary model of fractionation involving primocrysts and an interstitial trapped liquid. this interstitial liquid is now thought likely to have undergone significant in situ fractionation, including the preferential loss of at least one immiscible component. until we can quantify the amount of interstitial liquid and its fate, our understanding of the post-cumulus evolution of layered mafic intrusions will remain uncertain. one possible way forward is to characterise plagioclase overgrowths and their relationship to the volume of interstitial melt (maaløe 1976a; toplis et al. 2008, namur et al. 2014; nielsen et al. 2015; namur & humphreys 2018). 5.13 summary of solidification models fractional crystallisation is widely thought to have occurred during the solidification of the skaergaard magma chamber (wager & deer 1939; mcbirney 1989a, 1995), particularly as recorded in the major elements and the systematic variation in mineral compositions (e.g. figs 25, 26). perfect fractional crystallisation, modelled by rayleigh equations, however, fails to reproduce the observed variations specifically for the excluded trace elements (low d*). meanwhile the included trace and the major elements (high d*) are in many cases a better fit (figs 60, 62). the main obstacle to understand the solidification of the skaergaard intrusion is therefore, that this observation suggests a decoupling of the excluded and included trace elements, meaning that identical fractionation processes cannot model the two groups of trace elements. perfect fractional crystallisation is an end-member model that probably rarely operates in magma chambers where solidification is occurring from the topdown, bottom-up and sides-inward. in such a chamber, mush zones are present to a varying extent, and the efficiency of crystal-liquid separation is strongly controlled by the orientation of the solidification front and gravitational forces. under such variable conditions, the overall evolution would be delayed relative to that predicted for perfect fractional crystallisation. crystallisation is ultimately controlled by partition coefficients and phase equilibria governing crystallisation in a convecting and homogenising magma. the resultant solid substratum or crystal-melt mush is thus made up of minerals and melt in equilibrium with the main magma. in contrast, in situ (or bottom) crystallisation assumes crystallisation in a boundary mush layer or directly on a solid substratum, which is only in part equilibrated with the main magma and may thus be controlled by phase equilibria and bulk partitions that are different from those of the main magma chamber. the case for in situ fractionation of the skaergaard intrusion is preferred by nielsen et al. (2015) and nielsen (2016) based on the concentric form of the https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 114 of 139 geusbulletin.org intrusion (nielsen 2004), the evidence for the presence of abundant interstitial melt-mineral reactions (holness et al. 2011) and liquid immiscible droplets (jakobsen et al. 2011). however, for in situ crystallisation to be a viable possibility, the amount of required melt remaining in the mush must greatly exceed the predicted amounts (fig. 61; table 24). this indicates that either fractional crystallisation and compaction were the likely processes dominating the solidification or, as suggested by this study, that the predictions of the trapped-melt content are greatly underestimated (fig. 68; see also mcbirney 1995; tegner et al. 2009). the effects of liquid immiscibility on the terminal fractionation of the intrusion offer a mechanism by which to explain the decoupling of the included and excluded trace elements. by formulating simple models (figs 65, 66) that invoke migration and loss of granophyre melts formed from immiscibility initiated at the base of the uzb and likewise in the trapped interstitial melt, reasonable correspondence between perfect rayleigh fractionation and observed or inverted melt compositions can be obtained (fig. 67). this modelling works for at least some trace elements for which partition information is available and requires upward migration from the lz and uza into the main chamber as well as lateral migration. the main chamber at the uzb and uzc levels (the final 10%) is viewed as a zoned chamber with restricted convention or no convention, and it is suggested that granophyre melt is lost to the ubs and surroundings in the form of granophyre sills. liquid immiscibility near the end-stages of solidification may have resulted in the stagnation and possible stratification of the residual magma that suppressed efficient crystal-melt segregation, a process that had dominated the evolution of the main magmatic system up to that point. another consequence is that the skaergaard intrusion solidified with significantly higher amounts of trapped melts than predicted by tegner et al. (2009) and summarised in table 24. this would have resulted in an upper part (uzb, uzc) where the trapped-melt content might have reached high values approaching the sh horizon (fig. 68). https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 115 of 139 geusbulletin.org 6 concluding remarks a sound knowledge of the amount of trapped melt in the gabbro mush is crucial for understanding the solidification of the skaergaard magma chamber, including thermal and temporal cooling, the liquid line of descent, trace element partitioning and textural and mineralogical variability of the late-stage solidification of the gabbros. it is nevertheless apparent from the present study that such information is uncertain at best. the amount of trapped melt is calculated using excluded trace elements (like rb) analysed in the gabbros and normalised to the predicted concentration in the melt assuming perfect fractional or equilibrium crystallisation. only if the analysed gabbros represent the final gabbro mush and the crystallisation mode is sufficiently understood, will this method of calculation yield the correct result. tegner et al. (2009) assumed that the analysed samples represented the final mush gabbro and that the mode of crystallisation was approximated by rayleigh fractionation. the main drawback of these assumptions is that for extreme crystallisation with little remaining melt (f → 0), the trapped melt will always go to zero, irrespective of any actual observations to the contrary, such as the presence of interstitial granophyre. this is because the concentration of a highly excluded trace element will go to very high values when f goes to zero, assuming perfect fractional crystallisation. despite these shortcomings, particularly for the uzb and uzc, the procedure of tegner et al. (2009; fig. 61) has provided baseline calculations of the trapped-melt content used here and in other studies (e.g. mckenzie 2011; holness et al. 2017b). the mobility of a late-stage granophyric melt in the skaergaard mush has nevertheless, directly or indirectly, been acknowledged in the past (e.g. wager & brown 1967; naslund 1989; larsen & brooks 1994; jakobsen et  al. 2011) and has been built into summation modelling of the initial or final composition and the liquid descent line (tegner 1997; nielsen et al. 2009; andersen 2006). this has often been done without considering the implications of a missing granophyre or melanogranophyre component. a critical observation in the present study is that the excluded trace elements are decoupled from the included trace (and major) elements and thus cannot be explained by identical differentiation mechanisms (figs 62–64). a crystallisation model (figs 65, 66) illustrates the potential effects of liquid immiscibility in the trapped melt as well as in the overall magma chamber that is inferred to be initiated at the uza–uzb boundary or equivalent in the mush melt. using this model, it is possible to predict the evolution of both the included and the excluded elements for identical crystallisation modes (fig. 67). the requirement is that a granophyre component would have migrated upwards and laterally in the intrusion and, thus be partly lost from the sampled and analysed mush gabbros. these observations imply that the gabbro mush samples used here and in previous studies (e.g. tegner et al. 2009) may underestimate the actual trapped-melt component unless adjusted for these lateral and horizontal granophyre migrations. the differences between rubidium concentrations of the observed and the predicted modelling by perfect or imperfect fractional crystallisation without granophyre loss can be used to adjust the gabbro to its possible original composition (using c. 65 ppm rb in granophyre; table 18). the granophyre predicted to have been lost from the gabbros are shown in fig. 68a ranging from small amounts in the lza (c. 5%), raising fast to 25% in lzc, then further raising to 30% at the base of uzb, subsequently decreasing to 25% at the base of uzc, and thereafter dropping to low values (fig. 68a). the fundamental requirement for bottom solidification processes is the operation of a convecting and homogenising magma chamber concurrently with the growth of the solid mush along the margins and the expulsion of evolved melt, whether resulting from compaction or convective wall processes (wager et al. 1960; maaløe 1978; morse 1986; mcbirney 1995). it is the convection and homogenising of the main chamber that drives the systematic cryptic variation seen in the ls and ubs (figs 25–29). wager & brown (1967, p. 206) thought that the magma was essentially stationery during the formation of the so-called tranquil division of the mbs and subsequently became convective with the appearances of prominent banding. maaløe (1976a, 1987) placed the onset of large-scale convection in the chamber to about 35 m height in the cambridge drill core, approximately at the base of the lza. at this point, he observed a significant drop in plagioclase composition and change in zoning pattern. holness et al. (2015) also examined the cambridge drill core and suggested a pulsed filling of the lower parts (hz), followed by a rapid ballooning of the chamber commencing in lza, and by the onset of chamber-wide convention (cf. holness et al. 2007a). chamber-wide convection driven by cooling is believed to have persisted until the uza–uzb boundary. at this point, about 400 m of chamber remained in the central part of the intrusion (fig. 30), the melt composition may have reached an andesitic composition with about 62 wt% sio2 (fig. 44), and has seen a dramatic increase in viscosity to 7000 pa·s (for a dry si-enrichment https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 116 of 139 geusbulletin.org descent line; fig. 59). rhythmic layering reaches a maximum intensity in the uza, subsequently ceases in uzb, and is largely not recognisable in uzc (wager & deer 1939; wager & brown 1967; mcbirney & noyes 1979; irvine 1987). an effect of the homogeneous nature of uzb without prominent layering is the restricted modal and compositional variability observed in the bulk rock and primocryst of the samples from drill core 90-22 (figs 15, 25–27). holness et al. (2007b) observed a cessation in grain boundary adjustments in the uzb gabbro at about the same time as modal and grain-size layering became less prominent. they discussed these observations as the result of either a transition from convective to conductive heat transfer or to the cooling below the closure t for pyroxene-plagioclase grain boundary migration. namur & humphreys (2018) modelled the trace element contents in plagioclase and clinopyroxene to be broadly consistent with efficient fractional crystallisation until the lower part of uza (f ≈ 0.20). above this stratigraphic level, they proposed that fractional crystallisation became less efficient and was replaced predominantly by equilibrium crystallisation. our findings show that the uza–uzb boundary (c. 1060°c, f = 0.10) marks the onset of liquid immiscibility in the main chamber. this is supported by the modal variation (figs 15, 51) and the phase equilibria (fig. 47). based on a study of gabbro compositions, it is observed that the included and excluded trace elements are decoupled throughout most of the ls. only the included elements (thus also the major elements) can reasonably be modelled by fractional crystallisation (figs 62, 64). the apparent decoupling of included and excluded trace elements may arise from difficulties obtaining gabbro samples that represent the true final mush. the residual trapped melt in the mush cools in situ either isolated or by reactions with the primocrysts. it may thus reach saturation at 1060°c of two conjugate immiscible melts of which the basic melt is likely to remain in place and the silicic, or granophyric, melt may raise or otherwise migrate in the mush (fig. 59), possibly into the main chamber reservoir. the granophyre content lost from the gabbro mush increases with decreasing overall temperature. if the gabbro compositions are restored to their assumed residual mush compositions by adding a granophyre melt composition, it is possible to restrain the melt composition to perfect rayleigh fractionation into the uzb (fig. 66). this way, crystallisation can explain both the included and the excluded elements by fractional crystallisation in the lower part and by equilibrium in the upper part of the intrusion. the result is that the two trace element groups are no longer seen as being decoupled (fig. 67). namur & humphreys (2018) supports this conclusion. they used in situ trace element analyses of coexisting mush minerals without directly using the gabbro compositions, and reached the conclusion that the ls, until the lower part of uza, was controlled by fractional crystallisation after which they suggested an increasing component of equilibrium crystallisation. the adjusted trapped-melt content reaches high values in uz (fig. 68), which could significantly have changed the overall differentiation pattern and the behaviour of the mush. a smaller amount of melt would in the case of in situ fractionation have returned to the main chamber and consequently have dampened the fractionation effect and resulted in a better balancing of the included and the excluded elements (fig. 64). values of d* would have decreased (included elements) or increased (excluded elements) towards unity (→ 1) and have dampened the extreme modelled compositions for rayleigh fractionation (fig. 67). because a larger amount of melt was retained in the mush, another effect would have been that the gabbro column with stratigraphic height would have moved more rapidly towards low remaining melt content (f → 0). since the calculations of the adjusted trapped-melt content were based on rayleigh fractionation, it is not possible to further evaluate the relative contributions of perfect, imperfect, or in situ fractionation to the overall differentiation of the skaergaard intrusion. to gain better insight into the details of the differentiation process, independent estimates of the amount of trapped mush melt are required. most studies, including ours, would argue that fractional crystallisation controlled by in situ crystallisation and a major chamber homogenised by convection explains the observed differentiation in the ls (cf. mcbirney & noyes 1979; namur & humphreys 2018). based on modelling of a simplified basalt system (haplobasalt), maaløe (1976b) suggested that rayleigh fraction dominated, but argued for a component of imperfect fractionation (partial fractionation) to explain the delayed fractionation of plagioclase. nielsen et al. (2015) advocated for dominant in situ fractionation in the final stages of crystallisation and formation of the pge and gold deposit in what is commonly referred to as the triple group of the upper part of the mz. they have extended this idea to the entire intrusion (nielsen 2016) but have not provided substantial documentation and arguments for in situ fractionation as opposed to other modes of fractionation. higher volumes of interstitial mush melt would be present when reached the onset of liquid immiscibility at about 1060°c. consequentially, at this point the interstitial melt would still be present, perhaps forming an interconnected network that would better allow the buoyant migration of the silica-rich immiscible conjugate melt. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 117 of 139 geusbulletin.org the same observations may also provide insight into the local and large-scale metasomatic reactions proposed by mcbirney and coworkers (mcbirney 1989a, 1995, 2002; mcbirney & sonnenthal 1990; sonnenthal & mcbirney 1998). mcbirney (1996) writes that “[..] a buoyant liquid, rich in volatiles and incompatible elements, rose from the ls and permeated the upper part of the intrusion.” he attributed such an upward flux of reactive aqueous liquid or fluid in the ls mush to the observed pervasive infiltration metasomatism (mcbirney & sonnenthal 1990), as well as to the overall decoupling between excluded and included elements (mcbirney 1989a, 2002). specifically, sonnenthal (1992) interpreted evidence from cl/f ratios of apatite, biotite, and amphiboles in the uza to the exsolution of an aqueous cl-rich fluid (mcbirney 1995). changes in the apatite cl/f ratio was also observed in this study suggesting the exsolution of a cl-rich fluid at the uza–uzb boundary, followed by a second exsolution of a water-rich fluid at the uzb– uzc boundary (fig. 24b). mcbirney (2002) used a larger set of apatite analyses to place the initiation of the exsolution of a cl-rich fluid to the upper part of the mz. later work by pedersen et al. (2021), however, was unable to define the horizon where a potentially cl-rich fluid might have exsolved. the implications of the present study are that (1) compaction was restricted, (2) solidification occurred in a boundary-layer mush containing variable amounts of interstitial melt, (3) it was possible for a late-stage silica-rich conjugate immiscibility melt to migrate and raise from the mush while a fe-rich conjugate melt had limited ability to sink or was retained at the level of formation, (4) a prolonged upward migration of buoyant silicic melt into the main chamber was possible, and (5) final solidification of the mush package, after termination of immiscibility, was controlled by cooling to, or below, a closure temperature of c. 950°c. this scenario possesses some similarities to the suggestion by mcbirney (2002). complications arise from crystallisation of interstitial melt in pore spaces and the possibility of an upward migrating, reacting and crystallising melt (mcbirney 1998, 2002) that is not well understood. tegner et al. (2009) argued that compaction of the primocryst pile preceded the growth of interstitial trace element enriched minerals. some migration of interstitial basaltic melt from the trapped mush likely occurred (humphreys 2009; holness et al. 2011; holness 2018; namur et al. 2014), although such migration is believed to have had restricted impact. another factor at play may be compositional convection and exchange in a boundary zone of evolved mush melts with less-evolved magma chamber melts. this would result in constant composition of mush melt and thus an effective buffering of mineral compositions, eliminating trapped melt and interstitial crystallisation (wager et al. 1960; irvine 1980; morse 1986). tegner et al. (2009) and namur et al. (2014) have argued against such a possibility as a dominating factor for the skaergaard intrusion. they argue that given that very similar rock textures are universally and simultaneously developed in gabbro of the floor, walls, and roof of the intrusion, this is unlikely to be due to a gravitationally driven process of convection, but more likely to have been caused by compaction. additional processes are thus required to explain the observed excluded elemental deficiency compared to the predicted composition expected for perfect fractional crystallisation, including upward migration of a late-stage silicic melt from the intrusion due to compaction and density differences into the ubs and host rocks. mcbirney (2002), nevertheless, preferred upward elemental migration aided by an exsolved chlorine-rich, hydrous fluid from the mz and above that eventually escaped from the intrusion with its load of excluded elements. acknowledgements the icp-ms trace element analyses were done by eric brown and michelle gras at the interdisciplinary center for plasma mass  spectrometry, university of california, davis. the major element xrf analyses were done by sidsel grundvig at department of geoscience, aarhus university. the fluorine analyses were done by ole stecher while at university of copenhagen’s arctic station in qeqertarsuaq, disko, greenland. the electron microprobe analyses were done at the department of earth and planetary science, university of california, davis, and at the department of geoscience, aarhus university. jakob keiding previously analysed some of the uzc samples used here by electron microprobe in aarhus and completed gabbro density measurements. we are indebted to many more people who over the years have helped and enlightened our understanding of the skaergaard intrusion by always being willing to discuss even the most ephemeral subjects and to read our various nascent manuscripts. they include troels nielsen, kent brooks, gry barfod, rune larsen and harry esbensen. the final manuscript benefitted from critical comments from the editor jakob kløve keiding and from the reviewers ilya veksler and rais latypov. the bulletin editor catherine jex and her team greatly contributed to the final version of our manuscript. it is a special pleasure to be able to dedicate our work to alex mcbirney and neil irvine whose work on the skaergaard intrusion has been a constant inspiration. we must, however, take the full responsibility for any shortcomings in the present presentation and interpretations. additional information funding statement the danish lithosphere centre, funded by the danish national research foundation, supported our visits to the skaergaard intrusion in 1995 and 2000. the analytical program was supported by the us national science foundation under grants nsf-ear-0208075 and -1019887 (cel). the carslberg foundation and the independent research fund denmark supported ct. final stages of manuscript preparation was supported by danish national research foundation niels bohr professorship (26-123/8). author contributions all authors equally contributed to the conceptual ideas and formulation of research goals and planning and completion of the fieldwork. pt developed the methodology and models and wrote the original draft https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 118 of 139 geusbulletin.org and presentation of the result, while all authors contributed to the final review and editing. ct curated the field collected samples. cel acquired the main funding for the analytical aspects leading to this publication. competing interests the authors declare no conflicts of interests. additional files included in this manuscript are three appendices: appendix 1. history of exploration, appendix 2. sampling methods and reference profile construction and appendix 3. analytical methods. supplementary files are available to download at https://doi. org/10.22008/fk2/utfjp4. supplementary file s1: excel spreadsheet with summary of gabbro and mineral compositions. supplementary files s2: excel spreadsheets of all tables. supplementary files s3: forward modelled liquid lines of descent (tables s1, s2). supplementary files s4: plane-polarised microscope images of thin sections. supplementary files s5: classic sample location maps. supplementary files s6: geological maps. supplementary files s7: topographic maps. supplementary files s8: stereoscopic aerial colour photos (1:27 000). supplementary file s9: sample repositories. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ https://doi.org/10.22008/fk2/utfjp4 https://doi.org/10.22008/fk2/utfjp4 thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 119 of 139 geusbulletin.org references almeev, r.r., holtz, f., koepke, j., parat, f. & botcharnikov, r.e. 2007: the effect of h2o on olivine crystallization 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in felsic magmas: an assessment of the role of apatite. geochimica et cosmochimica acta 45, 2349–2358. https://doi. org/10.1016/0016-7037(81)90088-0 watson, e.b. & green, t.h. 1981: apatite/liquid partition coefficients for the rare earth elements and strontium. earth and planetary science letters 56, 405–421. https://doi.org/10.1016/0012-821x(81)90144-8 watts, griffis & mcouat ltd. 1991: 1990 skaergaard project, platinova/ corona concession, east greenland. exploration report, 55 pp. geological survey of denmark and greenland unpublished report, grf 20848. whitaker, m.l., nekvasil, h., lindsley, d.h. & difrancesco 2007: the role of pressure in producing compositional diversity in intraplate basaltic magmas. journal of petrology 48, 365–393. https://doi.org/10.1093/ petrology/egl063 williams, r.j. 1971: reaction constants in the system fe-mgo-sio2-o2: intensive parameters in the skaergaard intrusion, east greenland. american journal of science 271, 132–146. https://doi.org/10.2475/ ajs.271.2.132 wilson, j.r. & larsen, s.b. 1985: two-dimensional study of a layered intrusion – the hyllingen series, norway. geological magazine 122, 97–124. https://doi.org/10.1017/s0016756800031022 wilson, j.r. & sørensen, h.s. 1996: the fongen-hyllingen layered intrusive complex, norway. in: cawthorn, r.g. (ed.): developments in petrology 15, 303–329: elsevier. https://doi.org/10.1016/ s0167-2894(96)80011-x wilson, j.r., esbensen, k.h. & thy, p. 1981: igneous petrology of the synorogenic fongen-hyllingen layered basic complex, south-central scandinavian caledonides. journal of petrology 22, 584–627. https:// doi.org/10.1093/petrology/22.4.584 wotzlaw, j.-f., bindeman, i.n., schaltegger, u., brooks, c.k. & naslund, h.r. 2012: high-resolution insights into episodes of crystallization, hydrothermal alteration and remelting in the skaergaard intrusive complex. earth and planetary science letters 355–356, 199–212. https://doi.org/10.1016/j.epsl.2012.08.043 wright, j.b. 1961: solid-solution relationships in some titaniferous iron oxide ores of basic igneous rocks. mineralogical magazine 32, 778– 789. https://doi.org/10.1180/minmag.1961.032.253.04 wyllie, p.j. 1963: effects of the changes in slope occurring on liquidus and solidus paths in the system diopside-anorthite-albite. mineralogical society of america, special publication 1, 204–212. yang, h.-j., kinzler, r.j. & grove, t.l. 1996: experiments and models of anhydrous, basaltic olivine-plagioclase-augite saturated melts from 0.001 to 10 kbar. contributions to mineralogy and petrology 124, 1–18. https://doi.org/10.1007/s004100050169 yoder, h.s. & sahama, t.g. 1957: olivine x-ray determinative curve. american mineralogist 42, 475–491. yoder, h.s., tilley, c.e. & schairer, j.f. 1963: pyroxenes and associated minerals in the crust and mantle. pyroxene quadrilateral. carnegie institution washington, yearbook 62, 84–94. zeh, a., ovtcharova, m., wilson, a.h. & schaltegger, u. 2015: the bushveld complex was emplaced and cooled in less than one million years – results of zirconology, and geotectonic implications. earth and planetary science letters 418, 103–114. https://doi.org/10.1016/j.epsl.2015.02.035 https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ https://doi.org/10.1007/s00410-006-0127-y https://doi.org/10.1007/s00410-006-0127-y https://doi.org/10.1093/petrology/egm056 https://doi.org/10.1093/petrology/egm056 https://doi.org/10.1093/petrology/egn064 https://www.geochemsoc.org/files/3015/3608/9023/gnews_139_april_2009.pdf https://www.geochemsoc.org/files/3015/3608/9023/gnews_139_april_2009.pdf https://doi.org/10.2138/am.2010.3456 https://doi.org/10.2138/am.2010.3456 https://doi.org/10.1093/petrology/egh042 https://doi.org/10.1093/petrology/egh042 https://doi.org/10.1016/0016-7037(54)90026-5 https://doi.org/10.1007/s00410-018-1466-1 https://doi.org/10.1093/petrology/1.3.364 https://doi.org/10.1017/s0016756800060829 https://doi.org/10.1016/0016-7037(51)90016-6 https://doi.org/10.2113/gsecongeo.52.8.855 https://doi.org/10.1093/petrology/1.1.73 https://doi.org/10.1093/petrology/1.1.73 https://doi.org/10.1007/bf00371487 https://doi.org/10.1007/bf00371487 https://doi.org/10.1007/bf00374440 https://doi.org/10.1007/bf00374440 https://doi.org/10.1007/bf00375424 https://doi.org/10.1029/gl006i012p00937 https://doi.org/10.1029/gl006i012p00937 https://doi.org/10.1016/0016-7037(81)90088-0 https://doi.org/10.1016/0016-7037(81)90088-0 https://doi.org/10.1016/0012-821x(81)90144-8 https://doi.org/10.1093/petrology/egl063 https://doi.org/10.1093/petrology/egl063 https://doi.org/10.2475/ajs.271.2.132 https://doi.org/10.2475/ajs.271.2.132 https://doi.org/10.1017/s0016756800031022 https://doi.org/10.1016/s0167-2894(96)80011-x https://doi.org/10.1016/s0167-2894(96)80011-x https://doi.org/10.1093/petrology/22.4.584 https://doi.org/10.1093/petrology/22.4.584 https://doi.org/10.1016/j.epsl.2012.08.043 https://doi.org/10.1180/minmag.1961.032.253.04 https://doi.org/10.1007/s004100050169 https://doi.org/10.1016/j.epsl.2015.02.035 thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 129 of 139 geusbulletin.org appendix 1 history of exploration of the skaergaard intrusion the skaergaard intrusion was discovered in 1930 by l.r. wager as a member of the british arctic route expedition (chapman 1932; watkins 1932). the first aerial photos of the intrusion were made during this expedition and were later used by wager & deer (1939) for mapping purposes. wager returned in 1932 to the area as part of an ejnar mikkelsen expedition (mikkelsen 1933) exploring the district between ammassalik ø (tasiilaq) and scoresby sund (ittoqqortoormiit; wager 1934, 1935, 1947). wager returned in 1935 to the kangerlussuaq fjord (central east greenland), specifically to study the skaergaard intrusion, and overwintered in hjemstedbugt (homestead bay) on the intrusion with a small party, including his wife phyllis and w.a. deer (deer 1967a; brooks 1985, 1990; hargreaves 1991). the outcome of this overwintering was the seminal monograph on the skaergaard intrusion that appeared only three years later (wager & deer 1939). the interest was so strong that an unprecedented second amended version was issued in 1962. the years 1934–1972 saw several additional monographs published in the east greenland series of meddelelser om grønland based on results obtained during the british east greenland expeditions increasing the total published monographs to 10. the intrusion saw little activity during world war ii, except for the establishment of a weather station in 1945 on skærgårdshalvø, later transferred to aputiteeq. wager and deer returned in 1953 to the kangerlussuaq fjord to collect further samples from the skaergaard intrusion as well as for other mapping objectives in the area (wager 1954). new samples were needed, mainly due to rapid depletion of the original existing material from the numerous mineralogical and geochemical investigations that had since been conducted. wager incorporated this new information in a long chapter on the skaergaard intrusion included in a book on ‘layered igneous rocks’ written and edited by himself and g.m. brown (wager & brown 1967) that also included an updated version of the initial map from the 1939 monograph. this book was published in the us by w.h. freeman of san francisco in 1967 and the following year in 1968 by oliver & boyd of edinburgh, united kingdom. to investigate the unexposed basal parts of the intrusion (i.e. the hidden zone; hz), drilling was essential and preparation was underway in the 1960s. the untimely death of wager in late 1965, prior to the publication of the layered intrusion book, left the drilling plans to be concluded by w.a. deer, who returned to skaergaard the following year (deer 1967b). this drilling resulted in a total of 588 m of core collected from two sites penetrating gabbro (referred to as the cambridge drill cores). the only part of these drill cores that has been studied to some extent is core i, which penetrated 349 m into basal gabbros of the intrusion, including c. 150 m of unexposed stratigraphy (e.g., maaløe 1976; nwe 1976; holness et al. 2015). uk scientific interests in the skaergaard ceased with the cambridge drilling and the activities shifted to us groups mainly lead by a.r. mcbirney (oregon university), who organised several expeditions between 1971 and 1984 (e.g. naslund 1984; hoover 1989; mcbirney 1989a). this field work led to the compilation of the third and most recent geological map of the skaergaard intrusion in 1989 (mcbirney 1989b). the supplementary material includes scans of all the known geological maps. concurrently, c.k. brooks (university of copenhagen) organised expeditions to the kangerlussuaq area between 1972 and 1991, mainly focused on regional tectono-magmatic aspects, but also included studies of skaergaard on several occasions (larsen & brooks 1994; brooks 2018). t.n. irvine of the carnegie institution for science, washington, participated in several of these expeditions over the years and published extensively on the origin of the layering of the intrusion (e.g. irvine et al. 1998). d.k. bird from stanford university, california, was also part of these activities (bird et al. 1986). aerial photography by the danish geodetic institute was completed in 1975 eventually leading to a new, improved topographic map of the intrusion used by mcbirney and later investigators. see supplementary files s7 for the topographic maps and supplementary files s8 for high-resolution scans of the aerial photos used in this mapping. commercial exploration activities on skaergaard were initiated in 1986 with the discovery of a gold-palladium mineralisation (e.g. bird et al. 1991). several drill cores, mostly through the upper part of the intrusion, were completed by platinova resources ltd during 1989–1990 (watts, griffis & mcquat ltd 1991). some of these cores had been made available for research from an outdoor storage site near the intrusion (e.g. tegner 1997; andersen et al. 1998). subsequently, the commercial interest in the skaergaard mineralisation was transferred to skaergaard minerals corporation, who conducted further drilling and sampling that culminated in 2004 (hanghøj 2005). see the summary of sample repositories in the supplementary file s9 for the availability of skaergaard samples. despite that to date, several companies have been involved in the exploration activities on the skaergaard intrusion, which were still https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 130 of 139 geusbulletin.org ongoing at the time of writing. despite this, no attempt has been made so far to mine the deposit, mainly due to the high cost of mining in the arctic. the scientific interest expanded when parties from the danish lithosphere centre’s campaign on the east greenland rifted margin visited the intrusion in 1995 and 2000 (nielsen et al. 2001). several of the exploration drill cores from temporary outdoor storage in greenland were as a result transferred to permanent storage in copenhagen (nielsen et al. 2000). a unesco international geoscience programme field conference on site at the skaergaard intrusion was held in september 2001 organised by j.c.ø. andersen and c.k. brooks. following the conference, a ‘virtual’ skaergaard intrusion website has been maintained by j.c.ø. andersen (http:// www.skaergaard.org). brooks (1997, 2018) has published extensive and very useful accounts of the exploration of the skaergaard intrusion, including an annotated literature list with selected abstracts. most recently, groups from the university of aarhus (denmark), cambridge university (uk), and university of california davis (usa) continued work post 2000 on the skaergaard intrusion, both in laboratories and the field. part of these activities occurred concurrently with the renewed exploration drilling activities by skaergaard minerals corporation (hanghøj 2005). the present work is one of many outcomes of these recent efforts. wager & deer’s 1939 monograph has significantly shaped the theory of igneous petrology since it was first published. in 1939, the theory of magmatic differentiation and fractional crystallisation were being formulated and the common basalt liquid lines of descent discussed (wilson 1993; young 1998, 2003). n.l. bowen at the geophysical laboratory in washington dominated this debate and advocated a si-enrichment trend principally based on experimental observation (bowen 1928). the most prominent dissenting voice was c.n. fenner also at the geophysical laboratory that based his arguments for the existence of an fe-enrichment trend on field observations (e.g. fenner 1929). when wager & deer’s 1939 skaergaard study was published it landed squarely in the middle of this discussion by presenting a well documented example of what was seen as an fe-enrichment trend in an extreme differentiated pluton. it provided the foundation for understanding fractional crystallisation, convecting magma chambers and the origin of layering in igneous plutons. however, it is perhaps a tribute to the complexities of understanding the details of the skaergaard intrusion that the debate of si-enrichment vs. fe-enrichment trends (or bowen vs. fenner trends) still reverberate in the discussions and that this issue has not yet been satisfactory solved (hunter & sparks 1990). although our understanding today may have changed on some points, the foundation provided by wager & deer (1939) is still valid and still forms a central part of all igneous petrology textbooks (e.g. carmichael et al. 1974; hess 1989). references andersen, j.c.ø., rasmussen, h., nielsen, t.f.d. & rønsbo, j.g. 1998: the triple group and the platinova gold and palladium reefs in the skaergaard intrusion; stratigraphic and petrographic relations. economic geology 93, 488–509. https://doi.org/10.2113/gsecongeo.93.4.488 bird, d.k., rogers, r.d. & manning, c.e. 1986: mineralized fracture systems of the skaergaard intrusion, east greenland. meddelelser om grønland. geoscience 16, 68 pp. bird, d.k., brooks, c.k., gannicott, r.a. & turner, p.a. 1991: a gold-bearing horizon in the skaergaard intrusion, east greenland. economic geology 86, 1083–1092. https://doi.org/10.2113/gsecongeo.86.5.1083 bowen, n.l. 1928: the evolution of the igneous rocks. princeton: princeton university press. 332 pp. brooks, c.k. 1985: l. r. wager and the geology of east greenland. in: drake, e.t. & jordan, w.m. (eds): geologists and ideas 1, 237–250: geological society of america. https://doi.org/10.1130/ dnag-cent-v1.237 brooks, c.k. 1990: wager, laurence rickard. dictionary of scientific biography, supplement ii, 968–970. new york: charles schribner’s sons. brooks, c.k. 1997: the skaergaard intrusion: sixty years of petrological research. introduction and abstracts. danish lithosphere centre and institute of petrology, university of copenhagen, open file report. brooks, c.k. 2018: over eighty years at the core of petrological research: the skaergaard intrusion. the history of research, its environment and annotated bibliography. copenhagen: geological survey of denmark and greenland. carmichael, i.s.e., turner, f.j., & verhoogen, j. 1974: igneous petrology. mcgraw-hill, new york. 739 pp. chapman, f.s. 1932: northern lights: the official account of the british arctic air-route expedition 1930-31. london: chatto and windus, 304 pp. deer, w.a. 1967a: laurence rickard wager. 1904-1965. biographical memoirs of fellows of the royal society 13, 359–385. http://www.jstor. org/stable/769388 https://doi.org/10.1098/rsbm.1967.0019 deer, w.a. 1967b. east greenland geological expedition, 1966. polar record 13, 783–784. https://doi.org/10.1017/s0032247400058484 fenner, c.n. 1929: the crystallization of basalts. american journal of science s5-18(105), 225–253. https://doi.org/10.2475/ajs.s5-18.105.225 hanghøj k. 2005: report on exploration activities in 2004 on skaergaard license no. 2005/09. internal report, skaergaard minerals corp., 51 pp., 1 appendix: geochemical analyses, core recovery, drill hole survey, 175 pp. (in archive of the geological survey of denmark and greenland, report grf 21895, http://maps.greenmin.gl/ geusmap/?mapname=greenland_portal). hargreaves, j. 1991: l.r. wager: a life, 1904-1965. oxford: joshua associates. 140 pp. hess, p.c. 1989: origin of igneous rocks. cambridge: harvard university press. 344 pp. holness, m.b., tegner, c., namur, o. & pilbeam, l. 2015: the earliest history of the skaergaard magma chamber: a textural and geochemical study of the cambridge drill core. journal of petrology 56, 1199–1227. https://doi.org/10.1093/petrology/egv034 hoover, j.d. 1989: petrology of the marginal border series of the skaergaard intrusion. journal of petrology 30, 399–439. https://doi. org/10.1093/petrology/30.2.399 hunter, r.h. & sparks, r.s.j. 1990. the differentiation of the skaergaard intrusion. reply to a.r. mcbirney, h.r. naslund, s.a. morse, c.k. brooks & t.f.d. nielsen. contributions to mineralogy and petrology 104, 248–254. https://doi.org/10.1007/bf00306449 irvine, t.n., andersen, j.c.ø. & brooks, c.k. 1998: included blocks (and blocks within blocks) in the skaergaard intrusion: geological relations and the origin of rhythmic modally graded layers. geological society of america bulletin 110, 1398–1447. https://doi.org/10.1130/0016-7606 (1998)110<1398:ibabwb>2.3.co;2 https://doi.org/10.34194/geusb.v56.8327 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skaergaard intrusion, east greenland. journal of petrology 35, 1651–1679. https://doi.org/10.1093/petrology/35.6.1651 mcbirney, a.r. 1989a: the skaergaard layered series: 1. structure and average compositions. journal of petrology 30, 363–397. https://doi. org/10.1093/petrology/30.2.363 mcbirney, a.r. 1989b: geological map of the skaergaard intrusion, east greenland. university of oregon, eugene, oregon. mikkelsen, e. 1933: the scoresby sound committee’s 2nd east greenland expedition in 1932 to king christian ix×s land. report on the expedition. meddelelser on grønland 104(1), 71 pp. maaløe, s. 1976: the zoned plagioclase of the skaergaard intrusion, east greenland. journal of petrology 17, 398–419. https://doi.org/10.1093/ petrology/17.3.398 naslund, h.r. 1984: petrology of the upper border series of the skaergaard intrusion. journal of petrology 25, 185–212. https://doi. org/10.1093/petrology/25.1.185 nielsen, t.f., hansen, h., brooks, c.k., & lesher, c.e. 2001: the east greenland continental margin, the prinsen af wales bjerge and new skaergaard intrusion initiatives. geology of greenland survey bulletin 189, 83–98. https://doi.org/10.34194/ggub.v189.5162 nielsen, t.f.d. et al. 2000: retrieval of platinova drill cores: a new skaergaard initiative. 2000 fall meeting, san francisco, usa, 15–19 december 2000. eos, transactions, american geophysical union 81(48), f1366 only. nwe, y.y. 1976: electron-probe studies of the early pyroxenes and olivines from the skaergaard intrusion, east greenland. contributions to mineralogy and petrology 55, 105–126. https://doi.org/10.1007/ bf00372758 tegner, c. 1997: iron in plagioclase as a monitor of the differentiation of the skaergaard intrusion. contributions to mineralogy and petrology 128, 45–51. https://doi.org/10.1007/s004100050292 wager, l.r. 1934: geological investigations in east greenland: part i. general geology from angmagsalik to kap dalton. meddelelser om grønland 105(2). 46 pp. wager, l.r. 1935: geological 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https://doi.org/10.1007/bf00372758 https://doi.org/10.1007/s004100050292 https://doi.org/10.1017/s0032247400043503 https://tidsskrift.dk/geografisktidsskrift/article/view/47924 https://tidsskrift.dk/geografisktidsskrift/article/view/47924 https://doi.org/10.1144/gsjgs.150.4.0611 https://doi.org/10.1515/9780691187723 https://doi.org/10.1515/9780691187723 thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 132 of 139 geusbulletin.org appendix 2.1 platinova drill cores airlifted to copenhagen in 2000 drill core no. depth azimut inclination altimeter utm interval (m) (m) e n shipped (m) 90–10 535 300 –70 2 552374 7560758 0–535 90–18 1026 320 –70 59 553403 7560044 0–1026 90–22 1052 340 –70 306 554604 7561147 0–1052 90–23 843 35 –70 577 557884 7561597 218–569 90–23a 842 35 –70 577 557884 7561597 559–841 90–24 1072 0 –70 502 555066 7561739 0–1072 total 4318 appendix 2 sampling methods and reference profile construction the new reference profile is constructed using drill core samples through the uza and uzb and into the upper part of the mz and surface samples extending the coverage of the ls into the lza and the uzc, including the sh and the adjoining part of the ubs. between 1986 and 1990, platinova resources ltd. conducted an extensive mineral exploration program on the skaergaard intrusion focusing on the potential for precious metal deposits in the ls. a series of exploration diamond drillings were completed in 1989 and 1990 for which the drill cores were stored outside near a small aircraft landing field in sødalen. in 1999, the responsibility for this material was transferred to the geus. funding from the danish natural science research council and greenland bureau of mines and petroleum provided the opportunity to retrieve selected cores during the 2000 field season. the cores totalled 4308 m and were considered to be of greatest scientific importance (nielsen et al. 2000, 2001). these cores are now curated by the geological museum at the university of copenhagen, denmark (appendix a2.1). in this study, we only use samples from hole 90-22 that traverses the mz to uza boundary and penetrated below the triple group in the mz. hole 90-22 represents 1052 m of gabbros from which 51 representative samples were selected (appendix a2.2). surface samples were collected in four sub-profiles for optimal stratigraphic continuity. in this way lza, lzb, and lzc were collected on uttental plateau, lzc on kraemer ø, mz along the north side of forbindelsesgletscher, and uzc and sh on basistoppen. the position of each surface sample was recorded by gps and altimeter readings and plotted on a topographic map and aerial photos. the true stratigraphic distance between samples was established by correcting for the local dip of layering. a total of 85 gabbro surface samples were used in the construction of the combined reference profile. the surface sample columns were combined with dip-corrected core logs for hole 90-22 to construct a composite stratigraphic column that links the individual profiles and provides accurate locations of zone boundaries and thicknesses. the total stratigraphic section of the ls calculated in this way is 2165 m. the total number of samples used to construct the reference profile was 136 (85 surface and 51 drill core samples) giving an average sample interval of 16 m. it is noteworthy that significant differences exist in the relative thicknesses of some zones and subzones compared to those given by wager & deer (1939). for example, our location of the lz–mz and mz–uz boundaries are stratigraphically higher and lower, respectively, resulting in a significantly reduced thickness for mz relative to the stratigraphic column of wager & deer (1939). likewise, we find that the relative thicknesses of lzc and uza are greater than reported by wager & deer (1939). the present samples are collected essentially in the same profiles as those of wager & deer (1939). we therefore ascribe the differences in the thicknesses to improved positioning of samples and zone boundaries by gps positioning and continuous drill core. a summary of the samples used are given in table a2.2, listing the stratigraphic height in m of the reference profile together with the volume relations calculated by tegner et al. (2009) and based on the volume relations of nielsen (2004). the field locations of the samples used by wager & deer (1939) and wager & brown (1967) are based on the detailed manuscript sample location map maintained by r.l. wager at oxford university, uk. the map is handdrawn on two sheets of paper on which the topographic base map together with the main zone boundaries have been traced in pencil. the scale of the original map is identical to that used for the geological map in wager & brown (1967). a total of 4188 sample locations are include on two map sheets that are captioned ‘working https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 133 of 139 geusbulletin.org appendix 2.2 position and stratigraphic height of samples in the ls, skaergaard intrusion sample id a stratigraphic zone/ profile section rock type altitude c latitude d longitude d utm e f f t g position (m) b subzone (m) ° ‘ ° ‘ e n (°c) 458633 2169 ubs basistoppen gabbro 499 68 9.67 –31 40.03 555325 7561425 0.000 1012 458634 2168 ubs basistoppen gabbro 499 68 9.67 –31 40.03 555325 7561425 0.000 999 458635 2165 ubs basistoppen gabbro 494 68 9.66 –31 40.01 555343 7561409 0.000 955 458636 2165 uzc basistoppen gabbro 494 68 9.66 –31 40.01 555343 7561409 0.000 995 458637 2165 uzc basistoppen gabbro 493 68 9.66 –31 40.01 555343 7561409 0.000 993 458638 2165 uzc basistoppen gabbro 492 68 9.66 –31 40.01 555343 7561409 0.000 1006 458639 2165 uzc basistoppen gabbro 491 68 9.66 –31 40.01 555343 7561409 0.000 1005 458640 2164 uzc basistoppen gabbro 489 68 9.66 –31 40.01 555343 7561409 0.001 1001 458641 2164 uzc basistoppen gabbro 488 68 9.66 –31 40.01 555343 7561409 0.001 1008 458642 2163 uzc basistoppen melanogabbro 493 68 9.66 –31 40.01 555343 7561409 0.001 994 458643 2150 uzc basistoppen gabbro 484 68 9.67 –31 40.03 555325 7561425 0.002 997 458644 2147 uzc basistoppen gabbro 478 68 9.67 –31 40.08 555291 7561422 0.003 1006 458645 2144 uzc basistoppen gabbro 473 68 9.67 –31 40.11 555275 7561420 0.003 1003 458646 2144 uzc basistoppen gabbro 472 68 9.67 –31 40.11 555275 7561420 0.003 1007 458647 2141 uzc basistoppen gabbro 468 68 9.67 –31 40.18 555227 7561425 0.003 1008 0.7 2091 uzc core 90-22-0.7m gabbro 0.010 1017 458652 2089 uzc basistoppen gabbro 442 68 9.75 –31 40.74 554835 7561559 0.010 1020 10.6 2081 uzc core 90-22-10.6m gabbro 0.011 458653 2075 uzc basistoppen gabbro 423 68 9.75 –31 40.83 554774 7561558 0.012 1022 20.8 2071 uzc core 90-22-20.8m gabbro 0.013 1013 31.1 2061 uzc core 90-22-31.1m gabbro 0.014 40.9 2051 uzc core 90-22-40.9m gabbro 0.015 1022 46.2 2046 uzc core 90-22-46.2m gabbro 0.016 47.4 2044 uzc core 90-22-47.4m gabbro 0.016 1014 458654 2032 uzc basistoppen gabbro 401 68 9.79 –31 40.91 554711 7561631 0.018 1020 458657 2030 uzc basistoppen melanogabbro 391 68 9.77 –31 41.02 554638 7561607 0.018 1028 458658 2030 uzc basistoppen melanogabbro 391 68 9.77 –31 41.02 554638 7561607 0.018 1000 67.6 2024 uzb core 90-22-67.6m average gabbro 0.019 1025 87.7 2004 uzb core 90-22-87.7m average gabbro 0.022 1028 107.7 1984 uzb core 90-22-107.7m average gabbro 0.025 1024 129.6 1962 uzb core 90-22-129.6m average gabbro 0.029 1025 149.8 1942 uzb core 90-22-149.8m average gabbro 0.032 1030 170.4 1921 uzb core 90-22-170.4m average gabbro 0.036 1028 189.9 1902 uzb core 90-22-189.9m average gabbro 0.039 1030 210.9 1881 uzb core 90-22-210.9m average gabbro 0.043 1032 229.4 1862 uzb core 90-22-229.4m average gabbro 0.046 1029 249.2 1843 uzb core 90-22-249.2m average gabbro 0.050 1039 270.3 1822 uzb core 90-22-270.2m average gabbro 0.054 1034 290.8 1801 uzb core 90-22-290.8m average gabbro 0.058 310.2 1782 uzb core 90-22-310.2m average gabbro 0.062 1033 323.8 1768 uzb core 90-22-323.8m average gabbro 0.065 344.1 1748 uzb core 90-22-344.1m average gabbro 0.069 1035 364.1 1728 uzb core 90-22-364.1m average gabbro 0.073 1039 384.0 1708 uzb core 90-22-384.0m average gabbro 0.077 1041 401.3 1691 uzb core 90-22-401.3m average gabbro 0.081 421.0 1671 uzb core 90-22-421.0m average gabbro 0.086 1038 441.7 1650 uzb core 90-22-441.7m average gabbro 0.090 1043 461.8 1630 uzb core 90-22-461.8m average gabbro 0.095 1040 471.8 1620 uza core 90-22-471.8m average gabbro 0.097 481.8 1610 uza core 90-22-481.8m average gabbro 0.100 1044 491.8 1600 uza core 90-22-491.8m average gabbro 0.102 1047 530.9 1561 uza core 90-22-530.9m average gabbro 0.112 1045 555.5 1536 uza core 90-22-555.5m average gabbro 0.118 565.3 1527 uza core 90-22-565.3m average gabbro 0.120 1055 585.8 1506 uza core 90-22-585.8m average gabbro 0.126 1053 648.4 1443 uza core 90-22-648.4m average gabbro 0.143 1056 660.6 1431 uza core 90-22-660.6m average gabbro 0.146 689.1 1403 uza core 90-22-689.1m average gabbro 0.154 1058 729.2 1363 uza core 90-22-729.2m average gabbro 0.166 1062 749.2 1343 uza core 90-22-749.2m average gabbro 0.171 (continued) https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 134 of 139 geusbulletin.org appendix 2.2 (continued) position and stratigraphic height of samples in the ls, skaergaard intrusion sample id a stratigraphic zone/ profile section rock type altitude c latitude d longitude d utm e f f t g position (m) b subzone (m) ° ‘ ° ‘ e n (°c) 769.1 1323 uza core 90-22-768.1m average gabbro 0.177 1056 789.7 1302 uza core 90-22-789.7m average gabbro 0.184 1061 809.7 1282 uza core 90-22-809.7m average gabbro 0.190 1064 824.2 1268 uza core 90-22-824.2m average gabbro 0.194 844.0 1248 uza core 90-22-844.0m average gabbro 0.201 1062 854.5 1237 uza core 90-22-854.5m average gabbro 0.204 873.7 1218 uza core 90-22-873.7m average gabbro 0.210 1060 893.6 1198 uza core 90-22-893.6m average gabbro 0.217 903.8 1188 uza core 90-22-903.8m average gabbro 0.220 1064 914.2 1178 uza core 90-22-914.2m average gabbro 0.224 1056 941.8 1150 uza core 90-22-941.8m average gabbro 0.233 1066 458266 1131 mz pukugagryggen average gabbro 289 68 11.09 –31 40.44 554989 7564068 0.239 1068 458265 1098 mz pukugagryggen average gabbro 271 68 11.13 –31 40.48 554961 7564142 0.251 1064 458264 1060 mz pukugagryggen average gabbro 224 68 11.14 –31 40.52 554933 7564152 0.264 1071 458263 1055 mz pukugagryggen average gabbro 221 68 11.15 –31 40.59 554882 7564172 0.266 1064 458262 1053 mz pukugagryggen average gabbro 210 68 11.14 –31 40.63 554854 7564158 0.267 1073 458261 1032 mz pukugagryggen average gabbro 192 68 11.16 –31 40.65 554841 7564184 0.274 1070 458261a 1032 mz pukugagryggen average gabbro 192 68 11.16 –31 40.65 554841 7564184 0.274 458260 1024 mz pukugagryggen average gabbro 160 68 11.14 –31 40.78 554748 7564149 0.277 1073 458259 1010 mz pukugagryggen average gabbro 152 68 11.16 –31 40.80 554736 7564180 0.283 1075 458258 1009 mz pukugagryggen average gabbro 142 68 11.16 –31 40.85 554703 7564181 0.283 1073 458256 1003 mz pukugagryggen average gabbro 132 68 11.16 –31 40.87 554691 7564184 0.285 1078 458257 1003 mz pukugagryggen average gabbro 132 68 11.16 –31 40.87 554691 7564184 0.285 1075 458255 1002 mz pukugagryggen average gabbro 127 68 11.16 –31 40.89 554676 7564180 0.286 1075 458254 988 mz pukugagryggen average gabbro 117 68 11.18 –31 40.93 554647 7564219 0.291 1074 458253 984 mz pukugagryggen average gabbro 99 68 11.17 –31 41.01 554589 7564203 0.292 1079 458252 961 mz pukugagryggen average gabbro 84 68 11.19 –31 41.01 554591 7564233 0.301 1083 458251 953 mz pukugagryggen average gabbro 74 68 11.19 –31 41.11 554518 7564240 0.304 1076 458250 935 mz pukugagryggen average gabbro 63 68 11.21 –31 41.14 554497 7564283 0.311 1077 458249 930 mz pukugagryggen average gabbro 47 68 11.21 –31 41.18 554474 7564271 0.313 1081 458248 921 mz pukugagryggen average gabbro 36 68 11.22 –31 41.22 554446 7564293 0.316 1083 458247 920 mz pukugagryggen average gabbro 30 68 11.22 –31 41.24 554431 7564290 0.317 1065 458246 900 mz pukugagryggen average gabbro 1 68 11.24 –31 41.32 554370 7564323 0.325 1128 458289 890 mz kræmer ø average gabbro 75 68 11.24 –31 42.82 553334 7564306 0.329 1080 458288 877 lzc kræmer ø average gabbro 74 68 11.26 –31 42.86 553309 7564345 0.334 1081 458287 875 lzc kræmer ø average gabbro 95 68 11.30 –31 42.89 553287 7564409 0.335 1093 458286 847 lzc kræmer ø average gabbro 86 68 11.33 –31 42.86 553308 7564468 0.346 1086 458285 830 lzc kræmer ø average gabbro 87 68 11.35 –31 42.88 553294 7564510 0.353 1087 458284 816 lzc kræmer ø average gabbro 85 68 11.37 –31 42.86 553306 7564538 0.358 1088 458283 808 lzc kræmer ø average gabbro 95 68 11.39 –31 42.88 553287 7564577 0.362 1086 458281 798 lzc kræmer ø average gabbro 94 68 11.40 –31 42.89 553280 7564597 0.366 1086 458282 798 lzc kræmer ø average gabbro 94 68 11.40 –31 42.89 553279 7564597 0.366 1082 458280 784 lzc kræmer ø average gabbro 95 68 11.42 –31 42.91 553270 7564634 0.372 1085 458279 752 lzc kræmer ø average gabbro 84 68 11.44 –31 42.94 553247 7564680 0.385 1090 458210 742 lzc uttental plateau average gabbro 700 68 12.82 –31 40.63 554787 7567276 0.390 1089 458209 723 lzc uttental plateau average gabbro 658 68 12.77 –31 40.74 554711 7567187 0.398 1099 458278 710 lzb kræmer ø average gabbro 59 68 11.46 –31 42.90 553272 7564720 0.403 1093 458208 703 lzb uttental plateau average gabbro 637 68 12.77 –31 40.78 554684 7567177 0.406 1100 458277 703 lzb kræmer ø average gabbro 59 68 11.47 –31 42.93 553252 7564730 0.406 1089 458207 681 lzb uttental plateau average gabbro 62 68 11.50 –31 42.93 553248 7564790 0.416 1094 458276 681 lzb kræmer ø average gabbro 592 68 12.72 –31 40.88 554618 7567090 0.416 1094 458206 634 lzb uttental plateau average gabbro 542 68 12.71 –31 41.00 554537 7567062 0.437 1099 458205 592 lzb uttental plateau average gabbro 498 68 12.70 –31 41.11 554462 7567036 0.456 1102 458204 580 lzb uttental plateau average gabbro 481 68 12.69 –31 41.14 554440 7567017 0.462 1109 458203 558 lzb uttental plateau average gabbro 453 68 12.67 –31 41.20 554400 7566989 0.472 1109 458202 519 lzb uttental plateau average gabbro 406 68 12.65 –31 41.32 554315 7566952 0.491 1100 458201 488 lzb uttental plateau average gabbro 363 68 12.63 –31 41.40 554261 7566913 0.506 1107 458233 447 lzb uttental plateau average gabbro 331 68 12.64 –31 41.50 554196 7566921 0.526 1112 458232 424 lzb uttental plateau average gabbro 308 68 12.64 –31 41.51 554185 7566923 0.537 1109 458231 367 lzb uttental plateau average gabbro 282 68 12.67 –31 41.65 554092 7566971 0.566 1107 (continued) https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 135 of 139 geusbulletin.org locality sheet / l.r.w.’ wager’s manuscript locality maps are in the wager archives at oxford university museum of natural history (m104); uk. wager’s manuscript locality map is included in the supplementary files s5, in high-resolution pdf format. the field locations of the samples used by mcbirney (1989a) are available as manuscript locations on a copy of mcbirney’s (1989b) printed geological map. this is now in the archives of the american museum of natural history (amnh) and also included here as part of the supplementary files s5. references mcbirney, a.r. 1989a: the skaergaard layered series: 1. structure and average compositions. journal of petrology 30, 363–397. https://doi. org/10.1093/petrology/30.2.363 mcbirney, a.r. 1989b: geological map of the skaergaard intrusion, east greenland. university of oregon, eugene, oregon. nielsen, t.f.d. 2004: the shape and volume of the skaergaard intrusion, greenland: implications for mass balance and bulk composition. journal of petrology 45, 507–530. https://doi.org/10.1093/petrology/egg092 nielsen, t.f.d. et al. 2000: retrieval of platinova drill cores: a new skaergaard initiative. 2000 fall meeting, san francisco, usa, 15–19 december 2000. eos, transactions, american geophysical union 81(48), f1366 only. nielsen, t.f., hansen, h., brooks, c.k., & lesher, c.e. 2001: the east greenland continental margin, the prinsen af wales bjerge and new skaergaard intrusion initiatives. geology of greenland survey bulletin 189, 83–98. https://doi.org/10.34194/ggub.v189.5162 tegner, c., thy, p., holness, m.b., jakobsen, j.k. & lesher, c.e. 2009: differentiation and compaction in the skaergaard intrusion. journal of petrology 50, 813–840. https://doi.org/10.1093/petrology/egp020 thy, p., lesher, c.e. & tegner, c. 2013: further work on experimental plagioclase equilibria and the skaergaard liquidus temperature. american mineralogist 98, 1360–1367. https://doi.org/10.2138/am.2013.4044 wager, l.r. & brown, g.m., 1967: layered igneous rocks. san francisco: freeman. 589 pp. wager, l.r. & deer, w.a. 1939: geological investigations in east greenland: part iii. the petrology of the skaergaard intrusion, kangerdlugssuaq, east greenland. meddelelser om grønland 105(4). appendix 2.2 (continued) position and stratigraphic height of samples in the ls, skaergaard intrusion sample id a stratigraphic zone/ profile section rock type altitude c latitude d longitude d utm e f f t g position (m) b subzone (m) ° ‘ ° ‘ e n (°c) 458227 346 lzb uttental plateau average gabbro 266 68 12.67 –31 41.68 554070 7566978 0.576 1115 458226 294 lzb uttental plateau average gabbro 237 68 12.70 –31 41.74 554026 7567025 0.604 1121 458225 221 lzb uttental plateau average gabbro 155 68 12.67 –31 41.90 553916 7566984 0.643 1120 458224 177 lzb uttental plateau average gabbro 220 68 12.67 –31 42.13 553761 7566964 0.667 1124 458221 175 lzb uttental plateau average gabbro 246 68 12.65 –31 42.47 553526 7566931 0.668 1122 458220 173 lza uttental plateau average gabbro 250 68 12.64 –31 42.60 553434 7566903 0.669 1126 458219 161 lza uttental plateau average gabbro 261 68 12.66 –31 42.66 553390 7566941 0.676 1119 458218 159 lza uttental plateau average gabbro 261 68 12.64 –31 42.81 553292 7566911 0.677 1111 458217 137 lza uttental plateau average gabbro 258 68 12.66 –31 42.88 553239 7566940 0.689 1117 458216 125 lza uttental plateau average gabbro 290 68 12.66 –31 43.03 553137 7566945 0.696 1119 458215 115 lza uttental plateau average gabbro 287 68 12.66 –31 43.12 553077 7566942 0.701 1124 458214 107 lza uttental plateau average gabbro 284 68 12.65 –31 43.21 553011 7566928 0.706 1129 458245 96 lza uttental plateau average gabbro 275 68 12.65 –31 43.29 552960 7566918 0.712 458213 84 lza uttental plateau average gabbro 273 68 12.65 –31 43.32 552935 7566924 0.719 1127 458212 54 lza uttental plateau average gabbro 261 68 12.66 –31 43.39 552887 7566942 0.736 1119 458242 27 lza uttental plateau average gabbro 250 68 12.67 –31 43.46 552841 7566954 0.752 1146 458211 7 lza uttental plateau average gabbro 248 68 12.68 –31 43.52 552796 7566968 0.764 1124 a names of samples from drill core 90-22 are the measured depth in metres. samples with the prefix 4582 and 4586 are surface samples from the reference profile collected in 2000. b the baseline of stratigraphic height is the mbs–lza boundary at uttentals plateau. c altitute (m) estimated from altimeter and/or topographic map. d latitude and longitude measured by gps. e utm coordinates calculated from latitude and longitude readings. f f is the mass fraction of magma remaining in the chamber calculated using equation (1) of tegner et al. (2009). g temperature (°c) calculated from average an content of plagioclase using the equation of thy et al. (2013) (t (°c) = 895 + 3.8 * an (%) + 19 (°c). https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ https://doi.org/10.1093/petrology/30.2.363 https://doi.org/10.1093/petrology/30.2.363 https://doi.org/10.1093/petrology/egg092 https://doi.org/10.34194/ggub.v189.5162 https://doi.org/10.1093/petrology/egp020 https://doi.org/10.2138/am.2013.4044 thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 136 of 139 geusbulletin.org appendix 3 analytical methods sample preparation surface samples were selected without visible alteration and weighing 2–4 kg. a reference piece (0.5–1 kg) and a slab for thin section were cut from each sample. the reference pieces were stored in the collection of the natural history museum of denmark (university of copenhagen), for future studies. weathered surfaces were removed with a diamond saw from the working sample. the remaining material was then crushed into <2 cm pellets in a hydraulic steel piston press and split into two portions, each of 30–40 g. these splits were grounded to c.10 mm fine powder in a corundum shatter box. the first split was ground for 3–5 min to pre-contaminate the shatter box and discharged. hereafter, the shatter box was cleaned in ethanol and the second split was ground, also for 3–5 min to achieve the powder size that was then used for all geochemical analyses. for the drill core samples, 10–16 cm long core sections were split along the long axis into pieces constituting, respectively, 1/3 and 2/3 by volume. the small reference split was stored at the natural history museum of denmark (university of copenhagen). a slab for thin section was cut from the upper end of the working split and the remaining material (200–500 g) was crushed into <2 cm aggregate in a steel-jaw crusher, followed by splitting and grinding in a tungsten-carbide shatter box. x-ray fluorescence (xrf) analyses loi was determined by heating the powder in air in a muffle furnace at 950°c for 3 h. fused glasses were prepared by mixing 0.75 g ignited powder with 3.75 g of fluore-x65 hp (a commercial flux from socachim fine chemicals consisting of 66 wt% li2b4o7 and 34 wt% libo2) in a 30 ml 95pt-5au crucible. the crucible was transferred to a muffle furnace and the contents melted twice for 5 minutes at 1150°c, with swirling of the crucible between melting. after fusion, the melt was poured into a red-hot, 32 mm 95pt-5au mold and quenched with air to produce a flat glass disc. the glass disc was used for the major analysis. powder pellets were prepared by mixing thoroughly 6 g powder with 1.0 g phenol formaldehyde (british bakelite company; resin r0214). after mixing, the powder appendix 3.1 lower limit of detection, precision and accuracy of xrf analyses counting statistics east greenland basalt (geus 95358) 1σ sd lld ave. 1 sd precision reference accuracy (n = 52) (%rsd) a value b (%) c major elements (wt%) sio2 0.18 0.01 49.77 0.11 0.2 50.14 –0.7 tio2 0.033 2.00 0.01 0.7 2.03 –1.3 al2o3 0.18 0.01 14.39 0.04 0.2 14.67 –1.9 feotot 0.14 0.01 12.36 0.03 0.2 12.09 2.3 mno 0.004 0.001 0.20 0.00 0.0 0.18 8.4 mgo 0.08 0.01 6.96 0.03 0.4 6.81 2.2 cao 0.09 0.01 11.49 0.06 0.5 11.36 1.2 na2o 0.05 0.01 2.38 0.06 2.5 2.27 4.7 k2o 0.01 0.01 0.25 0.00 1.6 0.26 –4.3 p2o5 0.08 0.01 0.19 0.00 0.7 0.19 –2.3 total 99.99 100.00 trace elements (ppm) v 4 332 39 12 330 0.8 cr 4 1 159 13 8 187 –14.9 ni 1.2 0.6 78 8 10 83 –6.1 cu 1.7 0.6 172 22 13 168 2.4 zn 1.5 0.7 97 4 4 100 –3.1 rb 0.2 0.6 sr 2.6 0.5 206 3 1 205 0.3 y 0.5 0.8 30 2 5 30 –0.2 zr 3.1 0.5 114 2 2 113 0.3 nb 0.2 0.6 8 1 15 8 –6.3 ba 5 2 54 10 19 50 7.5 pb 0.2 0.9 a %rsd calculated as 100 × (sd / ave.). b major elements from geological survey of denmark and greenland (l.m. larsen personal communication 2002) and trace elements from oregon state university (o. stecher personal communication 2003). c accuracy calculated as 100 × ([ave. – reference] / reference). lld: lower limit of detection. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 137 of 139 geusbulletin.org was placed in a die and pressed in a hydraulic press for 5 min at 30 tons. the pellet was placed in an oven operating at 110°c. after 30 min the phenol formaldehyde was set, and the pellet was ready for x-ray trace element analysis. the major and the trace element analyses were performed on a panalytical pw2400 x-ray spectrometer using superq software. for the major elements a 3 kw rh-tube was used operating at 50 kv and 55 ma along with px-1 multilayer for na and mg, pe crystal for al and si, ge crystal for p, lif(200) crystal for k, ca and ti and lif(220) crystal for mn and fe. the detector was a gas flow proportional counter using p10 gas (10% methane in ar). for mn and fe, this detector was used in tandem with a sealed xe detector. for the trace elements ba, la, ce, v and cr, the rh tube was operated at 50 kv and 55 ma, for the rest of the trace elements the rh tube was operated at 60 kv and 45 ma. lif(200) crystal and 100m collimator were used throughout. the detector for ba and la was a gas flow proportional counter, for v, cr and ce this counter was used in tandem with the sealed xe counter. a scintillation counter was used for the rest of the trace elements. a total of 44 international silicate rock reference materials, with compositions ranging from basaltic to rhyolitic, were used for the calibrations (govindaraju 1994, 1995). for the major elements the fundamental parameter matrix correction model in the superq software was used. for ni, cu, zn, rb, sr, y, zr, nb, pb, th, and u, the mass attenuation corrections are based on measuring the intensity of the compton kβ line of rh. for ba, la, ce, v, and cr, determination of the major elements were included, and the fp matrix correction model was used. the analytical uncertainty was estimated by counting statistics as well as from the concurrent analyses of an internal reference standard (appendix a3.1) appendix 3.2 precision and accuracy of icp-ms analyses using the w-2 and bhvo-2 reference standards isotope w-2 bhvo-2 ave. (ppm) precision accepted uncertainty accuracy (%) c ave. (ppm) precision accepted uncertainty accuracy (%) c (n = 6) %rsd a values (ppm) b 1σ sd (n = 6) %rsd a values (ppm) b 1σ sd 45sc 36.1 2.2 36 1.1 –0.3 32.5 2.7 32 13 –1.5 51v 270.5 2.7 260 12 –4.0 317.3 2.5 317 11 –0.1 52cr 92.7 10.7 92 4.4 –0.7 291.7 4.4 280 19 –4.2 60ni 65.0 6.1 70 2.5 7.1 114.2 4.4 119 7 4.0 59co 42.8 6.3 43 2.1 0.5 44.8 3.9 45 3 0.4 68zn 79.1 4.7 80 2 1.1 92.8 3.2 103 6 9.9 85rb 21.3 3.0 21 1.1 –1.3 9.5 2.9 9.8 1 2.6 88sr 201.1 5.1 190 3.0 –5.8 402.3 1.8 389 23 –3.4 89y 23.7 4.8 23 1.6 –2.9 27.8 1.7 26 2 –7.0 90zr 90.5 2.9 100 2 9.5 176.3 3.1 172 11 –2.5 93nb 7.5 3.9 7.9 5.6 18.3 4.2 18 2 –1.7 138ba 172.6 1.3 170 11 –1.3 130.3 1.7 130 13 –0.2 139la 10.6 1.5 10 0.59 –5.9 14.7 1.4 15 1 2.3 140ce 23.1 1.6 23 1.5 –0.6 36.3 1.2 38 2 4.4 141pr 3.0 1.9 0.0 5.2 1.3 0.0 146nd 13.2 2.1 13 1 –1.4 24.0 1.8 25 1.8 4.0 147sm 3.3 2.9 3.3 0.13 0.8 5.9 1.7 6.2 0.4 4.6 151eu 1.1 3.4 1.0 0.06 –11.1 2.0 1.7 0.0 157gd 3.7 3.4 0.0 5.9 2.4 6.3 0.2 7.1 159tb 0.6 3.1 0.63 2.7 0.9 2.2 0.9 –0.5 163dy 3.8 3.8 3.6 0.8 –5.4 5.2 1.5 0.0 165ho 0.8 2.8 0.0 1.0 1.6 1.04 0.04 8.5 167er 2.2 3.8 2.5 12.6 2.4 1.6 0.0 169tm 0.3 6.0 0.38 12.2 0.3 4.4 0.0 174yb 2.0 4.5 2.1 0.2 2.7 1.9 1.9 2.0 0.2 4.5 175lu 0.3 4.5 0.33 7.9 0.3 2.6 0.28 0.01 4.6 178hf 2.3 3.6 2.60 0.18 9.8 4.3 2.2 4.1 0.3 –4.7 184w 0.3 7.5 0.0 0.2 3.2 0.0 181ta 0.5 9.0 0.5 9.8 1.2 6.3 1.4 16.4 208pb 7.6 1.5 9.3 17.7 1.9 10.9 0.0 232th 2.3 5.3 2.4 0.1 5.6 1.2 0.9 1.2 0.3 3.2 238u 0.5 5.4 0.53 –1.1 0.4 3.2 a %rsd calculated as 100 × (sd / ave.). b data source: http://minerals.cr.usgs.gov/geo_chem_stand/index.html italic numbers are information values. c accuracy calculated as 100 × ((ave. – accepted) / accepted). w-2: usgs geochemical reference material w-2 (diabase). bhvo-2: usgs geochemical reference material (hawaiian basalt). https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ http://minerals.cr.usgs.gov/geo_chem_stand/index.html thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 138 of 139 geusbulletin.org feo was determined by decomposing the sample in a teflon crucible using a mixture of hydrofluoric and sulfuric acid at boiling point. the crucible and contents were plunged into boric acid solution. the ferrous iron liberated was titrated with potassium dichromate using barium diphenylamine sulfonate as indicator. icp-ms analyses transitional elements and ree were determined on whole-rock samples by icp-ms at the interdisciplinary center of plasma mass spectrometry at university of california, davis. a portion of the homogenised powdered samples was dried in an oven overnight. nearly 100 mg of each powder was weighed and placed in savellex beakers with a hf:hno3 solution. this closed vessel digestion procedure follows the method by jenner et al. (1990). the presence of refractory minerals in a few samples required microwave digestion. an agilent 7500 quadrapole icp-ms was used with certified standards bir-1, bhvo-1, bcr-2, and be-n for constructing calibration curved. drift was monitored by spiking each sample with an internal standard solution containing fixed concentrations of cs, ga, and in. daily tuning ensured the highest sensitivity and lowest variability (%rsd > 3%) using a 10 ppb solution that contains li, cr, y, ce, and tl. the production of ionised oxides (masses 156/140) and doubly charged ions (masses 70/140) were reduced to less than 1% and c. 2%, respectively. precision and accuracy are evaluated based on multiple analyses of reference standards w-2 and bhvo-2 (appendix a3.2). fluorine analyses fluorine was analysed using the ion selective electrode method of stecher (1998). the precision is believed to be less than 3–6% with a lower limit of detection of 15 ppm. the accuracy was evaluated by stecher (1983) to be within 5% of international reference rock samples. density determinations the bulk-rock specific density of the surface and drillcore samples was determined by weighing the dry and clean samples in air and in water. the precision of the density measurements is estimated at ± 1–2%. electron microprobe analyses mineral compositions were determined at either university of california at davis, usa, or at university of aarhus, denmark. the analyses at university of california at davis were done using cameca sx-50 or sx-100 electron microprobes with an acceleration voltage of 15 kv, a beam current of 10 na, and counting times between 10 and 30 s. all mineral phases were analysed using natural minerals as standards for calibration and wavelength dispersive spectrometers. a beam diameter of 1–2 µm was used to analyse most minerals, except that pyroxenes were analysed with a 20–30 µm broad beam to minimise the effects of exsolution. the analyses at the university of aarhus were done with a jeol 8600 superprobe equipped with three wavelength dispersive spectrometers and one energy-dispersive spectrometer. nearly all analyses were performed using a broad beam (c. 20 µm) operated at 20 kv and 10 na current, using synthetic and natural standards for calibration and zaf correction procedures. for pyroxenes, si, al, fe, mg, and ca were measured by energy-dispersive spectrometry with a 200 s counting time, and ti, mn and, na by wavelength dispersive spectrometry with a 40 s counting time. for olivine, si, fe, and mg were measured by energy-dispersive spectrometry with a 120 s counting time and mn and ni were determined by wavelength dispersive spectrometry using a 40 s counting appendix 3.3 laboratory comparison of electron microprobe analyses (batberg clinopyroxene) ucd au recommended average 1σ sd average 1σ sd (bb1989.286) (n = 35) (n = 45) sio2 54.20 0.28 54.96 0.19 54.52 tio2 0.11 0.02 0.10 0.02 0.09 al2o3 0.38 0.03 0.31 0.02 0.27 feo 1.34 0.07 1.29 0.09 1.13 mno 0.03 0.02 0.03 0.02 0.03 mgo 17.74 0.17 17.38 0.18 17.68 cao 24.53 0.24 24.73 0.17 24.21 na2o 0.52 0.04 0.43 0.14 0.45 cr2o3 0.67 0.05 0.63 0.05 (0.65) total 99.52 99.86 99.03 ucd: university of california at davis, usa, cameca microprobe. au: university of aarhus, denmark, jeol superprobe. bb1989.286: recommended value by xrf on bulk batberg clinopyroxene bb1989.286. https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ thy et al. 2023: geus bulletin 56. 8327. https://doi.org/10.34194/geusb.v56.8327 139 of 139 geusbulletin.org time. for plagioclase, si, al, ca, and na were determined by energy-dispersive spectrometry with a 120 s counting time, whereas fe and k were determined by wavelength dispersive spectrometry with 40 s counting time. when possible, a total of three points were analysed in the cores of each of three grains for each mineral (olivine, plagioclase, clinopyroxene, orthopyroxene, ilmenite, magnetite, apatite). the results are reported as averages and standard deviations for the specific thin section. analyses with compositional anomalies, suggesting impurities, and with oxide sums lower than 98.5% or higher than 101.5% having generally been excluded from the average values. an internal pyroxene standard was analysed at both laboratories concurrently as a control and measure of analytical precision and accuracy (appendix a3.3). least-squares mixing calculations modal proportions on a weight basis of the constituent minerals were calculated by weighted, leastsquares, linear approximations (bryan et al. 1969) using the mineral and the bulk-gabbro compositions. the weighting factors were 1.0 for all oxides, except for sio2 and al2o3 which were assigned factors of 0.4 and 0.5, respectively. the following mineral phases were included in the calculations throughout the ls: olivine, plagioclase, clinopyroxene, orthopyroxene, ilmenite, magnetite, and apatite. orthoclase and quarts were also included for the uzc. only acceptable positive solutions to the least squares approximations were compiled with typical sums of the squares of the residuals below 0.10. references bryan, w.b., finger, l.w. & chayes, f. 1969: estimating proportions in petrographic mixing equations by least-squares approximation. science 163(3870), 926–927. https://doi.org/10.1126/ science.163.3870.926 govindaraju, k. 1994: 1994 compilation of working values and sample description for 383 geostandards. geostandards newsletter 18(s1), 1–158. https://doi.org/10.1046/j.1365-2494.1998.53202081.x-i1 govindaraju, k. 1995: 1995 working values with confidence limits for twenty-six crpg, anrt and iwg-git geostandards. geostandards newsletter 19(s1), 1–32. https://doi.org/10.1111/j.1751-908x.1995. tb00164.x jenner, g.a., longerich, h.p., jackson, s.e. & fryer, b.j. 1990: icp-ms — a powerful tool for high-precision trace-element analysis in earth sciences: evidence from analysis of selected u.s.g.s. reference samples. chemical geology 83(1), 133–148. https://doi. org/10.1016/0009-2541(90)90145-w stecher, o. 1983: fluorine in twenty-two international reference rock samples and a compilation of fluorine values for the usgs reference samples. geostandards newsletter 7(2), 283–287. https://doi. org/10.1111/j.1751-908x.1983.tb00384.x stecher, o. 1998: fluorine geochemistry in volcanic rock series: examples from iceland and jan mayen. geochimica et cosmochimica acta 62(18), 3117–3130. https://doi.org/10.1016/s0016-7037(98)00210-5 https://doi.org/10.34194/geusb.v56.8327 http://www.geusbulletin.org/ https://doi.org/10.1126/science.163.3870.926 https://doi.org/10.1126/science.163.3870.926 https://doi.org/https http://doi.org/10.1046/j.1365-2494.1998.53202081.x-i1 https://doi.org/https http://doi.org/10.1111/j.1751-908x.1995.tb00164.x http://doi.org/10.1111/j.1751-908x.1995.tb00164.x https://doi.org/https http://doi.org/10.1016/0009-2541(90)90145-w http://doi.org/10.1016/0009-2541(90)90145-w https://doi.org/https http://doi.org/10.1111/j.1751-908x.1983.tb00384.x http://doi.org/10.1111/j.1751-908x.1983.tb00384.x https://doi.org/https http://doi.org/10.1016/s0016-7037(98)00210-5 petrology of the skaergaard layered series 1 introduction 2 the skaergaard intrusion 3 methods 4 stratigraphy of the layered series (ls) 4.1 petrography and zonal division 4.1.1 lower zone (lz) 4.1.1.1 lower (sub)zone a (lza) 4.1.1.2 lower (sub)zone b (lzb) 4.1.1.3 lower (sub)zone c (lzc) 4.1.2 middle zone (mz) 4.1.3 upper zone (uz) 4.1.3.1 upper (sub)zone a (uza) 4.1.3.2 upper (sub)zone b (uzb) 4.1.3.3 upper (sub)zone c (uzc) 4.1.4 liquid immiscibility 4.2 mineral mode and variation in gabbro density 4.3 mineral chemistry 4.3.1 olivine 4.3.2 plagioclase 4.3.3 clinopyroxene 4.3.4 ferrobustamite 4.3.5 orthopyroxene 4.3.6 feti oxides 4.3.7 apatite 4.3.8 biotite 4.4 cryptic mineral variations 4.4.1 olivine 4.4.2 plagioclase 4.4.3 pyroxenes 4.4.4 feti oxides 4.4.5 relative cryptic changes 4.4.6 co-variation of silicate minerals 4.5 bulk-rock compositions 4.5.1 major elements 4.5.2 trace elements 4.5.2.1 included trace elements 4.5.2.2 excluded trace elements 4.5.3 co-variation of mineral modes and trace element concentrations 5 discussion 5.1 liquidus temperatures 5.2 oxygen fugacity (fo2) 5.3 lithostatic pressure 5.4 end-stage melt compositions 5.5 liquid summation and the liquid line of descent 5.6 forward modelling of major oxides 5.6.1 crystallisation conditions and parental melt compositions 5.6.2 experimental foundation 5.6.3 fractional crystallisation model 5.6.4 results of forward modelling 5.7 additional phase-equilibria constraints 5.7.1 projection schemes 5.7.2 crystallisation order and zone divisions 5.7.3 clinopyroxene crystallisation 5.7.4 feti oxide crystallisation 5.7.5 orthopyroxene crystallisation 5.7.6 constraints on trapped-melt compositions 5.7.7 field of liquid immiscibility 5.7.8 apatite crystallisation 5.7.9 bulk-rock compositions and cotectic relations 5.8 iron-magnesium exchange equilibria 5.9 magmatic volatiles 5.10 kinematic and thermal constraints 5.11 forward modelling of trace elements 5.11.1 calculation methods 5.11.2 mineral-liquid partition coefficients (d) 5.11.3 bulk partition coefficients (d*) 5.11.4 trapped melt content 5.11.5 perfect fractional crystallisation 5.11.6 imperfect fractional crystallisation 5.11.7 in situ boundary-layer crystallisation 5.12 role of liquid immiscibility 5.12.1 evidence for liquid immiscibility 5.12.2 liquid immiscibility and crystallisation in the residual chamber 5.12.3 crystallisation and immiscibility in the trapped melt 5.12.4 predicting the effects of liquid immiscibility 5.12.5 observed, predicted and ideal liquid lines of descent 5.12.6 effects of granophyre migration on calculated trapped melt content 5.13 summary of solidification models 6 concluding remarks acknowledgements additional information funding statement author contributions competing interests additional files references appendix 2 sampling methods and reference profile construction references appendix 2 sampling methods and reference profile construction references appendix 3 analytical methods sample preparation x-ray fluorescence (xrf) analyses icp-ms analyses fluorine analyses density determinations electron microprobe analyses least-squares mixing calculations references figures fig. 1 north atlantic igneous province as outlined by sea-floor spreading ridges, major fracture zon fig. 2 summary of the skaergaard intrusion and its layered series (ls). a: schematic representation fig. 3 location of samples and the drill core used in the present study. a: geological map after mcb fig. 4 summary of calculated modal mineralogy of the examined gabbros displayed in the triangular (c fig. 5 representative petrography of the lz gabbros (samples 458287, 875 m; 458205, 592 m; 458211, 7 fig. 6 representative pyroxene intergrowths and late crystallising mineralogy and textures in lz gab fig. 7 representative petrography of the mz gabbros (samples 458251, 953 m and 458250, 935 m). shown fig. 8 representative pyroxene intergrowths and late crystallising mineralogy and textures in the mz fig. 9 antithetic relationship between the calculated modal contents of olivine and orthopyroxene sh fig. 10 representative petrography of uz gabbros (samples 458653, 2075 m; 90-22-344.1, 1748 m and 90 fig. 11 representative pyroxene intergrowths and late crystallising mineralogy and textures in uza a fig. 12 mineralogy and textures of ferrohedenbergite and inverted ferrobustamite intergrowths in a u fig. 13 elemental-mapped electron kα images (si, na, k, al) of interstitial granophyre patches in uz fig. 14 textures of melanogranophyre from uzc (458643, 2150 m) illustrated by a bse density image. t fig. 15 modal variation of the ls (fractions as wt%) as a function of stratigraphic height (m). calc fig. 16 olivine mineral chemistry showing mn as a function of mg as cations per formula unit (pfu) n fig. 17 composition of olivine coronas (fo mol%, see fig. 6) as a function of average primocryst com fig. 18 plagioclase mineral chemistry showing k and fe as a function of ca as cations per formula un fig. 19 the pyroxene quadrilateral diagram for the skaergaard pyroxenes on a molecular basis. a: nom fig. 20 variation diagrams as a function of mg content for selected cations of augite-ferrohedenberg fig. 21 variation diagrams as a function of mg, ca, and al content for elected cations and ratios of fig. 22 ferric iron estimate for the high-ca pyroxenes using the charge balance equation fe3+ = aliv fig. 23 percentage of the main modelled non-quadrilateral molecular components of the pyroxene solid fig. 24 compositions of minor mineral components of the ls. a: ternary fe2+-ti-fe3+ diagram showing fig. 25 cryptic variation of olivine as a function of stratigraphic height (m) in the ls. shown are fig. 26 cryptic variation in plagioclase as a function of stratigraphic height (m) in the ls. shown fig. 27 cryptic variation in pyroxenes (augite) as a function of stratigraphic height (m) in the ls fig. 28 cryptic variation in pyroxenes (augite) as a function of stratigraphic height (m) in the ls fig. 29 cryptic variation in ilmenite (blue) and magnetite (black) as a function of stratigraphic he fig. 30 derivative cryptic variation in the uzb and uzc for coexisting olivine (as fo mol%), plagioc fig. 31 mg/(mg + fetotal) content in coexisting mafic minerals (olivine, clinopyroxene, and orthopyr fig. 32 mg/(mg + fetotal) content in coexisting olivine (ol) and orthopyroxene (opx; %) vs. mg/(mg + fig. 33 major elements of the ls gabbros, shown as oxides (wt%) vs. stratigraphic height (m). blue h fig. 34 transition trace elements (sc, v, cr, co, ni, cu and zn; ppm) as a function of stratigraphic fig. 35 selected alkali metal (rb), alkaline earth (sr, ba), high field strength (y, zr, nb, hf), ra fig. 36 selected rare earth element (lan and ybn) concentrations (ppm) and sm/yb ratio normalised to fig. 37 selected trace elements (rb and la) concentrations (ppm) as a function of the trapped melt c fig. 38 rare earth elements (ree) of uz normalised to normal morb (log scale; sun & mcdonough, 1989) fig. 39 summary of calculated equilibrium temperatures (t; °c) using singleand two-phase thermomet fig. 40 summary of subsolidus oxygen fugacity (log fo2) vs. temperature (t) for coexisting magnetite fig. 41 subsolidus temperature (t, °c) and oxygen fugacity (fo2) as a function of stratigraphic heig fig. 42 calculations of the liquid line of descent and initial melt composition shown on feo* vs. si fig. 43 summary of the one-atmosphere melting experiments on skaergaard-related dykes of thy et al. fig. 44 forward modelling of the liquid line of descent and the cryptic mineral variation for the pr fig. 45 perspective view of the lz-mz gabbros in the quaternary olivine (ol) quartz (q) plagiocl fig. 46 projections from quartz (q) of the quaternary ol-q-pl-di on to the triangular di-pl-ol diagr fig. 47 projections from plagioclase (pl) of the quaternary ol-q-pl-di on the triangular di-ol-q dia fig. 48 pyroxene variation as a function of ls zones and subzones in the plagioclase (pl) projection fig. 49 normative fraction of diopside (di) of the modelled liquid line of descent (normative di/(di fig. 50 summary interpretations of projections from plagioclase (pl) of the quaternary ol-q-pl-di on fig. 51 variation in the ratio between modal olivine and clinopyroxene (molar weight fractions) as a fig. 52 interpretation of liquid immiscibility. projections from plagioclase (pl) and diopside (di) fig. 53 interpretation of the temperature (t; °c) of onset of liquid immiscibility as a function of fig. 54 modal content of apatite (wt%) as a function of modal olivine (wt%). curve is visual estimat fig. 55 the p2o5 concentrations (wt%) of melt saturated in apatite as a function of temperature (t; fig. 56 calculated partitioning kd of fe and mg for three different pairs of pyroxenes and olivine ( fig. 57 co-variation of fe/mg ratios for observed orthopyroxenes and olivine (table 3, 5 and 6) toge fig. 58 pyroxene quaternary diagram with phase equilibria (thin black lines) and temperature contour fig. 59 density (g·cm-3) and viscosity (pa·s) as a function of stratigraphic height (m) in the ls. a fig. 60 selected trace elements (sc, v, rb, ba, yb and u) from gabbros and modelled melt as a functi fig. 61 modelled trapped melt (%) as a function of stratigraphic height (m) in the ls. model uses d* fig. 62 inverted melt concentrations of rb, sc and yb (ppm) compared to variable rayleigh fractions fig. 63 modelled melt concentrations (ppm) of rb, sc and yb for imperfect fractional crystallisation fig. 64 modelled melt concentrations (ppm) for in situ crystallisation of rb, sc, and yb as a functi fig. 65 schematic illustration of the proposed late history of the terminal skaergaard magma chamber fig. 66 liquid immiscibility mass balance models for uzb and uzc. a: loss of granophyre from the int fig. 67 immiscibility models of rb, sc and ba (ppm) shown as a function of stratigraphic height (m) fig. 68 calculation of fraction of trapped melt (ftm) as a function of stratigraphic height (m) in t tables table 1 average modal makeup of zones and subzones of the ls (wt%) table 2 summary of calculated modes without trapped melt table 3 average olivine compositions table 4 average plagioclase compositions table 5 average clinopyroxene compositions table 6 pearson product moment correlation coefficient (r) matrices for non-quadrilateral and quadri table 7 average ferrobustamite compositions per sample table 8 average orthopyroxene compositions table 9 representative coexisting magnetite and ilmenite compositions per sample table 10 representative apatite compositions table 11 individual and average mica compositions table 12 summary of cryptic primocryst compositions table 13 summary of gabbro major element compositions table 14 average zone and subzone major and trace element compositions table 15 summary of gabbro trace element compositions table 16 summary of gabbro ree compositions table 17 pearson product moment correlation coefficient (r) matrices for mineral modes and selected table 18 summary of average skaergaard granophyric compositions compared to modelled compositions table 19 summary of calculated parental melt compositions for the skaergaard intrusion using the sum table 20 suggested initial melt compositions for the skaergaard intrusion (wt%) table 21 modelled silica enrichment melt compositions table 22 modelled iron enrichment melt compositions table 23 summary of selected single-element and bulk-element partition coefficients table 24 summary of preferred modes used for the trace element modelling including trapped melt table 25 variables used for modelling of liquid immiscibility appendix 2.1 platinova drill cores airlifted to copenhagen in 2000 appendix 2.2 position and stratigraphic height of samples in the ls, skaergaard intrusion appendix 3.1 lower limit of detection, precision and accuracy of xrf analyses appendix 3.2 precision and accuracy of icp-ms analyses using the w-2 and bhvo-2 reference standards appendix 3.3 laboratory comparison of electron microprobe analyses (batberg clinopyroxene) geological survey of denmark and greenland bulletin 21, 2010 geological survey of denmark and greenland bulletin 21 · 2010 exploration history and place names of northern east greenland anthony k. higgins geological survey of denmark and greenland ministry of climate and energy geological survey of denmark and greenland bulletin 21 keywords exploration history, northern east greenland, place names, lauge koch’s geological expeditions, caledonides. cover illustration ättestupan, the 1300 m high cliff on the north side of kejser franz joseph fjord discovered and so named by a.g. nathorst in 1899. frontispiece: facing page map of greenland by egede (1818), illustrating the incorrect assumption that the norse settlements of greenland were located in south-west and south-east greenland. many of the localities named in the icelandic sagas are placed on this map at imaginary sites on the unknown east coast of greenland. the map is from the second english edition of hans egede’s ‘description of greenland’, a slightly modified version of the first english edition published in 1741. chief editor of this series: adam a. garde editorial board of this series: john a. korstgård, department of earth sciences, university of aarhus; minik rosing, geological museum, university of copenhagen; finn surlyk, department of geography and geology, university of copenhagen scientific editor of this volume: adam a. garde editorial secretaries: jane holst and esben w. glendal referees: ian stone (uk) and christopher jacob ries (dk) illustrations: eva melskens maps: margareta christoffersen digital photographic work: benny m. schark layout and graphic production: annabeth andersen geodetic advice: willy lehmann weng printers: rosendahls · schultz grafisk a/s, albertslund, denmark manuscript received: 22 april 2010 final version approved: 1 july 2010 printed: 21 december 2010. minor factual corrections made may 2011 in the www.geus.dk version. issn 1604-8156 isbn 978-87-7871-292-9 velux fonden supported publication of this bulletin (see acknowledgements). citation of the name of this series it is recommended that the name of this series is cited in full, viz. geological survey of denmark and greenland bulletin. if abbreviation of this volume is necessary, the following form is suggested: geol. surv. den. green. bull. 21, 368 pp. available from geological survey of denmark and greenland (geus) øster voldgade 10, dk-1350 copenhagen k, denmark phone: +45 38 14 20 00, fax: +45 38 14 20 50, e-mail: geus@geus.dk or at www.geus.dk/publications/bull © de nationale geologiske undersøgelser for danmark og grønland (geus), 2010 for the full text of the geus copyright clause, please refer to www.geus.dk/publications/bull abstract . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . geographical limits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . exploration and discovery . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . scope of place names – approved / unapproved. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . official place names in greenland . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . exploration history of northern east greenland . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . pioneer exploration and discovery: c. 2500 bc – 1912 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . commercial activities, early mountaineering, geological mapping: 1919–1960 . . . . . . . . . modern scientific investigations, adventure and sporting expeditions 1961–2008. . . . . . . catalogue of place names in northern east greenland . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . glossary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . administrative organisation of greenland . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . abbreviations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . geographical terms . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . references. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . maps 1–3 (end of volume) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . maps 4 & 5 (folded in pockets) 5 7 8 9 11 12 13 17 18 30 65 117 350 350 350 350 351 365 contents 4 5 higgins, a.k. 2010: exploration history and place names of northern east greenland. geological survey of denmark and greenland bulletin 21, 368 pp. + 2 maps. the first recorded landing by europeans on the coast of northern east greenland (north of 69°n) was that of william scoresby jr., a british whaler, in 1822. this volume includes a chronological summary of the pioneer 19th century exploration voyages made by british, danish, norwegian, swedish, french and german expeditions – all of whom reported that the region had previously been occupied by the inuit or eskimo; also included are brief outlines of the increasing number of government and privately sponsored expeditions throughout the 20th century, whose objectives included cartography, geology, zoology, botany, trapping and the ascent of the highest mountain summits. in 1934 the place name committee for greenland was established, the tasks of which included a review of all place names hitherto recorded on published maps of greenland, their formal adoption in danicised form, and the approval or rejection of new name proposals. in northern east greenland, by far the largest numbers of new place names were those proposed by scientists associated with lauge koch’s geological expeditions that lasted from 1926 until 1958. this volume records the location and origin of more than 3000 officially approved place names as well as about 2650 unapproved names. the author’s interest in the exploration history and place names of northern east greenland started in 1968, when the geological survey of greenland initiated a major five-year geological mapping programme in the scoresby sund region. systematic compilation of names began about 1970, initially with the names given by william scoresby jr., and subsequently broadened in scope to include the names proposed by all expeditions to northern east greenland. the author has participated in 16 summer mapping expeditions with the survey to northern east greenland. publication of this volume represents the culmination of a lifetime working in the arctic. abstract authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark e-mail: akh@geus.dk 6 7 the place names of northern east greenland, be tween 69° and 81°21´n, provide a vivid record of the exploration of one of the least accessible parts of greenland. this region includes the eastern part of the north-east greenland national park, the largest national park in the world, and extends southwards beyond the national park limits to include the east greenland settlement of scoresbysund / illoqqor toor miut (ittoqqortoormiit) that was founded in 1925. illoqqortoormiut is the official spelling, while ittoq qortoormiit is the east greenlandic dialect spelling used by the inhabitants. all of east greenland was formerly occupied by inuit (eskimo) cultures, whose house ruins are found throughout the region, but none of the names the inuit used have survived. the region was re-discovered by whalers in the early 19th century and mainly explored by european expeditions of british, danish, norwegian, swedish, french and german origin. an increasing number of expeditions with varied objectives visited the region throughout the 20th century. after formation of the place name committee for greenland (stednavneudvalget) in 1934, the names used on all existing published maps were systematically reviewed and with few exceptions approved in danicised form. more than 190 place names used by the greenlandic inhabitants of scoresbysund / illoq qor toormiut (ittoqqortoormiit) since 1925 were re corded in 1955. in this volume the term ‘inuit’ is used in references to the former eskimo residents of northern east greenland, whereas ‘greenlandic’ or ‘green landers’ is used in respect of the present day in ha bitants. a very large number of place names were proposed by geologists and other scientists associated with lauge koch’s expeditions between 1926 and 1958, and reflect to some extent the diverse nationalities of the participants; names were given after persons, towns or geographical locations of danish, swiss, finnish, french, swedish and british origin. one of the principal reasons that so many names proposed by lauge koch’s scientists have been approved is that the journal meddelelser om grønland (published by the videnskabelige kommission for grønland: the scientific commission for greenland) insisted that only officially approved place names could be used in their publications. up until the 1960s meddelelser om grønland was the preferred publication for rou tine geological and other scientific descriptions, be cause it had the resources to produce well-illustrated accounts that were often accompanied by coloured folding maps. the author’s interest in east greenland’s exploration history and place names was stimulated during his first summer in east greenland in 1968. this visit to the scoresby sund region was undertaken with the geological survey of greenland (ggu – subsequently the geological survey of denmark and greenland, geus). between 1968 and 1998 the survey mapped geologically the entire region of northern east green land between 69° and 81°21´n, and the author participated in a total of 16 summer expeditions to this region with the survey. in about 1970 the author began compilation of the place names used on the various discovery and exploration expeditions that have visited east greenland since 1822, and this work has continued until publication of this volume. from 1990 onwards participants in the ggu/geus regional mapping expe di tions were supplied with collections and explanations of official place names relevant to the region of study (higgins 1990, 1994a, 1997). it is these preliminary col lec tions of place names that form the basis for the present work. a variety of publications on place names used in greenland exist, for example on the names that can be traced to the norse settlements in south-west and west greenland (vebæk 1966), and those that relate to the dutch whale-hunting period in west green land (bobé 1915, 1921; rosendahl 1974). the only attempt at a regional account on the origin of place names is dan laursen’s ‘the place names of north greenland’ (laursen 1972). laursen’s work has close similarities with this volume in that the great majority of place names listed relate to european and north american voyages of exploration. however, the pre sentation is somewhat verbose in giving very detailed information on each of the expeditions that visited north greenland, and the place name section is not presented in strict alphabetical order. when developing a style of presentation for this volume, i have followed in many respects that of the ‘dictionary of alaska place names’ by donald j. orth (1967) that introduction lists a very large number of place names in a compact and, for the reader, an informative and easily understood way. the monumental volume by orheim et al. (2003), ‘the place names of svalbard’, lists the more than 8000 currently approved names for that region, but has only a summary section describing the exploration of svalbard. the individual name entries are presented in a very summary way, such that it is not always obvious which expedition or person is responsible for the name. note that throughout this volume officially ap proved names are given in ordinary type, and in the place name catalogue in bold type. unapproved or unofficial names are always given in italics. the names of ships are given emphasis by use of capitals. geographical limits the traditional divisions of greenland are illustrated in fig. 1. east greenland (‘østgrønland’ in danish, ‘tunu’ in greenlandic) comprises the entire eastfacing coast from lindenow fjord / kanger lus suat siaq at 60°30´n to nordostrundingen at 81°21´n. the boundary between east greenland and north greenland (‘nordgrønland’ in danish, ‘avannaarsua’ in greenlandic) follows the sw–ne-trending watershed in kronprins christian land. this official boundary between north greenland and east green land is followed in this volume. east greenland can be conveniently divided into northern and southern re gions at c. 69°n, where a high ice cap and a long inhospitable coast have hindered migration of both the inuit and land animals. this natural boundary at 69°n has been adopted in this volume as the southern limit of ‘northern east greenland’. in 1976 the geological survey of greenland (ggu) introduced regional subdivisions of all of greenland that were considered more appropriate and useful for geological descriptions (fig. 2; escher & watt 1976). for northern east greenland the sub divisions chosen essentially follow the informal usage of lauge koch’s expeditions and other workers, with the notable exception of the boundary between ‘north’ and ‘east’ greenland, that is placed at an artificial limit of latitude 79°30´n rather than following the official boundary (fig. 1). the survey subdivisions thus somewhat illogically place the northernmost segment of the east-facing coast of northern east greenland in an enlarged ‘north greenland’. these survey subdivisions were first used extensively in the survey’s volume on the ‘geology of greenland’ (escher & watt 1976). the subdivisions were slightly amended by ghisler (1990), mainly to bring the offshore divisions into line with the onshore divisions. while these revised subdivisions have no formal official status they have been very widely used in geological publications for the past 30 years. indeed, the widespread usage particularly of the english term ‘north-east’ greenland has led to the assumption that there is an equivalent danish term for this part of east greenland, such that ‘nordøst grønland’ is commonly encountered in danish scientific publications and even in the formal title of ‘nordøstgrønlands nationalpark’ (north-east green land national park). 8 fig. 1. the three official divisions of greenland: tunu – østgrønland (east green land), avannaarsua – nordgrønland (north greenland), kitaa – vest grønland (west greenland). kmj: kap morris jesup; pl: peary land; kcl: kronprins christian land; dll: dronning louise land; r: rigny bjerg; g: gunnbjørn fjeld; kf: kap farvel. exploration and discovery a detailed summary of all significant expeditions to northern east greenland makes up the section on the exploration history of northern east greenland (see page 17). in this section the main phases of exploration are briefly outlined. maps 1–5 at the end of this volume give the most important place names used in northern east greenland. the former indigenous inhabitants of northern east greenland have left abundant evidence of their presence in the form of house ruins and tent rings. the inuit (eskimo) cultures can be related to several waves of immigration, of which the last few survivors of the thule culture in northern east greenland were probably the group of 12 encountered by douglas clav ering on clavering ø (74°15´n) in 1823 (clave ring 1830). the earliest names still applied to east greenland geographical features are those found in the icelandic sagas, but these were mainly given for distant high mountains used as landmarks when sailing to the norse settlements of south-west and west green land (østerbygden and vesterbygden; see frontis piece) from about ad 1000; these settlements gradually declined during the little ice age that followed, with the last certain contacts with europe about 1410. records of sightings of northern east greenland were few until henry hudson’s voyage in 1607 that observed hold with hope at 73°30´n and reported abundant whales in the waters near spitsbergen in the north atlantic. whalers of many nations flocked to spitsbergen after 1612, and when whales became scarce there about 1630 they began to sail to east green land waters in search of new hunting grounds. for more than 200 years, however, the coast of northern east greenland, protected by a wide belt of pack ice, was widely considered inaccessible. it was not until 1822 that the british whaler william scoresby jr. made the first recorded landings around the mouth of scoresby sund (70°15´n). from the mid-1800s onwards there were numerous visits to east green land waters by whalers, and notably norwegian sealers, who approached the land to supplement their catch with walruses and muskoxen. the german explorer karl koldewey made an attempt to reach the north pole via east greenland in 1869, but his ship germania only reached as far north as germania havn (74°32´n) where it was forced to overwinter. during the autumn and the spring of 1870, sledge journeys were sent northwards as far as 77°n, and the region from 74° to 77°n was mapped in outline for the first time. the next major mapping expedition was the danish expedition led by carl ryder in 1891–1892, that overwintered at hekla havn on danmark ø in the inner scoresby sund re gion (70°–72°n). the system of fjords was explored by boat and on sledge journeys. in 1899, a swedish expedition led by a.g. nathorst visited the kong oscar fjord region in a search for traces of salomon andrée’s balloon expedition that had vanished in 1897 during an attempt to reach the north pole. during the summer of 1899, nathorst explored kej ser franz joseph fjord and the network of fjords 9 canada station nord ellesmere island peary land western eastern central daneborg mestersvig pituffik / thule air base melville bugt svartenhuk halvø disko disko bugt upernavik nuuk aasiaat kangerlussuaqsisimiut tasiilaq / ammassalik scoresby sund blossevil le k ys t ivittuut qaqortoq kap farvel nare s s trai t central east southern east north east iceland north west north cent ral west southern west south east south west south arctic ocean north atlantic ocean inland ice 400 km 72° 76° 80° 28°36°44°52° 68° 64° 60° fig. 2. the subdivisions of greenland as used by ggu/geus. the geological survey’s subdivisions were based on unofficial usage in published reports by geologists, botanists, zoologists and other scientists. although not officially recognised these divisions have been very widely used in geological publications since 1976. centred on kong oscar fjord (72°–74°n); the expedition surveyor, per dusén, carried out an epic pro gramme of mapping. the 1906–08 danmark-ekspe ditionen was the largest and most ambitious of early danish expeditions, whose aims were to survey the large unknown region north of 77°n and to link up with the explorations of the american robert e. peary in north greenland. their success cost the lives of three members, of whom only the body of one has been found. norwegian activities entered a new phase with the first deliberate overwintering of a fox-trapping expedition in 1908–09. this was the start of the norwe gian–danish trapper era that was to last until 1960. a series of expeditions from both nations overwintered at hunting stations with networks of small hunting huts surrounding them, trapping foxes and occasional wolves for their skins. the expansion of their relative trapping terrains led to the trappers becoming involved in the norwegian-danish dispute over the sovereignty of east greenland that was settled in den mark’s favour at the international court of jus tice in the hague in 1933. during world war ii da nish and norwegian hunters co-operated as mem bers of nordøstgrønlands slædepatrulje (forerunner of the present sirius sledge patrol). trapping was resumed after the war but was only sustained with government subsidies, and when subsidies were suspended, falling skin prices led to the effective cessation of hunting in 1960. the full story of the trapping era is related in fascinating detail by peter schmidt mikkelsen (2008). members of the pioneer exploration voyages made occasional ascents of significant mountains, not ably julius payer during karl koldewey’s 1869–70 expedition, but voyages aimed primarily at climbs of the highest known mountains in northern east green land began with the british cambridge expedition led by j.m. wordie that travelled to east green land in 1926 aboard the heimland. the professed objectives in cluded surveying and archaeology, but included a reconnaissance of a route to the 2940 m summit of petermann bjerg, first seen by karl kolde wey’s expedition in 1870. wordie’s second expedition in 1929 was rewarded by its successful ascent. an italian expedition led by leonardo bonzi had sailed to east greenland with a small climbing expedition in 1934 intending to make an attempt on the watkins bjerge, the range of high summits south of scoresby sund at around 69°n. frustrated by ice conditions they explored the then unknown mountains of volquaart boon kyst (c. 70°n). it was another british expedition led by augustine courtauld and lawrence r. wager that made the first ascent of gunnbjørn fjeld (hvit serk) in august 1935; at 3694 m this is the highest peak in the watkins bjerge and the highest summit in greenland. these early climbing expeditions were all reliant on boats for transport. the competing interests of danish and norwe gian trappers led to signing of a treaty on east green land (østgrønlandstraktaten) in 1924 that allowed both nations to hunt, fish and carry out scientific investigations, but made no decision on sovereignty. however, the treaty specifically allowed denmark to establish a colony in the scoresby sund region, and this proposal was brought to fruition thanks to the influence and initiative of ejnar mik kelsen. in 1925 the greenlandic settlement of scores bysund / illoq qor toormiut (ittoqqortoormiit) was esta blished with the arrival of 70 greenlandic settlers, mainly from ammassalik / tasiilaq. this act and the series of geo logical expeditions initiated by lauge koch in 1926 were part of a strategy to expand danish influence in northern east greenland that eventually led to recognition of danish sovereignty over all of greenland. lauge koch’s 1926–27 expedition was followed by summer expeditions in 1929 and 1930 and then 1931– 34 treårsekspeditionen (the three-year expedition), the largest and most comprehensive expedition hitherto sent to east greenland by denmark, and also led by lauge koch. the danish geodætisk institut (geo detic institute) was an integral part of this expedition, and initiated a long-running programme of surveying leading to publication of 1:250 000 scale topographic maps. treårsekspe di tionen was succeeded by the socalled two-year expedition 1936–38, but the outbreak of world war ii led to a halt in scientific activities. lauge koch’s expeditions continued from 1947, with an almost entirely geological focus, until the annual grants for field work were abruptly suspended after the 1958 season. significant activities by other nations included the seven voyages to the scoresby sund region by jeanbaptiste charcot in his three-mast barque pourquoi pas?, that included the setting up of the french inter national polar year station 1932–33 in scoresbysund / illoqqortoormiut (ittoqqortoormiit) and the four voy ages by the american louise a. boyd with the vesle kari that visited most of the northern east greenland fjords undertaking photography and surveying. in 1952 an airport was constructed west of mesters vig, subsequently known as mestersvig (one-word), in connection with exploitation of the lead deposits 10 discovered by lauge koch’s expeditions nearby. the excellent 1800 m gravel airstrip has provided easy access to east greenland for aircraft, and between 1954 and 1985 about 200 scientific and sports expeditions made use of the airstrip facility to reach east greenland (mestersvig was partly replaced in 1985 by the new airport built at constable pynt). from the mid-1950s onwards climbing expeditions paid particular attention to the high mountains of the stauning alper that could be reached either by walking in, or by using small rubber boats for transport westwards along the coast (bennet 1972). in 1974 the national park in north-east greenland was established and in 1988 was expanded westwards across north green land. at present it is the largest national park in the world. the sirius sledge patrol, whose primary pur pose is to patrol the uninhabited regions of northern east and north greenland, also act as wardens in the national park; visitors can only enter the national park with permits issued by the greenland authorities. in 1967 the geological survey of greenland sent a small reconnaissance expedition to the scoresby sund region. this was a precursor to the major re gional geological mapping programme that was to prepare 1:500 000 scale geological maps of northern east green land over a period of 30 years (henriksen & higgins 2008). between 1978 and 1987 super wideangle, vertical, aerial photographs were taken covering all of greenland, and a new network of fixed survey stations was established by the geodætisk institut (gi, geodetic institute, now part of kort & matrikelstyrelsen – kms). kms has produced a new topographic database for all of greenland at a nominal scale of 1:250 000 based on a combination of digitised existing published 1:250 000 scale map sheets with new maps drawn of previously unmapped areas (http://en.nunagis.gl). in connection with their re gional geological mapping programmes, ggu/geus in co-operation with kms has prepared topographic maps on a 1:100 000 scale for almost all of northern east greenland. the sirius sledge patrol began to use twin otter aircraft for transport of personnel to and from dane borg in 1977, and from 1978 gi and ggu/geus expeditions to east and north greenland also made extensive use of twin otter aircraft, whose short takeoff and landing (stol) capabilities are ideal for transport of equipment and personnel between base camps and ‘unprepared’ natural landing strips (usually river terraces; p.s. mikkelsen 2006). twin otter aircraft chartered from iceland were also extensively used to supply the dye stations (part of the amer ican distant early warning radar system) on the inland ice, and to support the ice-drilling operations at various locations on the inland ice. the icelandic twin otter pilots thus achieved considerable experience in landing on snow and ice surfaces, and these skills have since been utilised by climbing expeditions to provide easy access to the high mountain ranges at around 69°n and other areas. climbing expeditions have also been transported by twin otter to less precipitous areas otherwise difficult to reach, and many of these expeditions appear to have had as their sole objective the ascent of unclimbed peaks that they can then name after themselves or members of their families. however, no unofficial names given by climbers have been recognised by the place name committee for greenland since about 1960. in many remote nunatak areas the summits climbed may be only a few hundred metres above the surrounding glacier surfaces on which the twin otter aircraft landed. claims by such expeditions to have made ‘30 first ascents’ are not unusual. scope of place names – approved / unapproved the bulk of this volume comprises a catalogue of approved and unapproved place names, arranged al phabetically, that have been used on maps and in publications for localities in northern east greenland (see page 117). there are more than 3000 officially recognised place names in the region 69°–81°21´n, that is to say names that have been approved by the place name com mittee for greenland in copenhagen (sted navne udvalget) established in 1934. in 1979 green land achie ved home rule (hjemmestyre), and in january 1984 the responsibility for place names in greenland was transferred to grønlands sprognævn in nuuk, the present nunat aqqinik aalajangiisartut / grønlands stednavnenævn. a review of the work of the place name committee for greenland from 1934 to de cember 1983 is the subject of the section that follows below: official place names in greenland (p. 13). several hundred place names that appeared on the 1:200 000 and 1:100 000 scale norwegian maps of parts of northern east greenland are also listed, although only a small number were approved, the great major11 ity being rejected by the then place name committee as being politically motivated, i.e. given to support norwegian arguments for claims to sovereignty over parts of east greenland. the detailed account of norwegian and danish trapping activities in northern east greenland by p.s. mikkelsen (1994, 2008) illustrates all their hunting stations and hunting huts, as well as the names and alternative names by which they are known. all these names receive brief mention here. unapproved names used on published maps by scientists of j.-b. charcot’s expeditions (1925–36), louise a. boyd’s expeditions (1932, 1935, 1948) and in the 1968 edition of ‘den grønlandske lods’ (this volume, published in danish, is ‘the greenland pilot for east greenland’) are also included. up to 1960 many of the names given by climbing expeditions to peaks in the stauning alper were ap proved in danicised form, but the proposals for an increasing number of foreign-sounding names led to adoption of a more critical attitude to approval of names by the place name committee for greenland. the existence of detailed topographic maps of the stauning alper has allowed identification of the positions of virtually all summits climbed up to 2008, and a special map on a 1:150 000 scale giving both ap proved and unapproved names applied to features in this region accompanies this volume as map 5. the large numbers of scientific, tourist and climbing expeditions that have visited, and continue to visit east greenland have inevitably led to the naming of geographical features. only a selection of unapproved names used for significant reference localities receive mention here. in general, any names given after living persons, or used for minor peaks or variations of climbers’ routes, are not included. the most important source of information for this volume has been a near complete set of the minutes of the former place name committee for greenland; these include the documentation submitted to the committee and its various sub-committees in considering place name proposals. this material was kindly lent to the author by henry w. bjørn of the then geodetic institute (geodætisk institute, now incorporated into kort& matrikelstyrelsen). acknowledgements peter schmidt mikkelsen (rønde, denmark) has kind ly allowed me to add to this volume the many variations of hut names used by trappers, and to quote the gps latitudes and longitudes he has determined for all the hunting stations and hunting huts in east greenland. this data, and the histories of the individual stations and huts, is taken from the english edition of his account of danish and norwegian trapping activities (p.s. mikkelsen 2008). jan løve (skagen, denmark) has for many years independently compiled data on place names used by expeditions to east greenland, backing up his compilations with studies of published and unpublished diaries and other original material in various danish archives. he has freely allowed me to make use of his deductions and conclusions with respect to specific names, thus correcting many of my errors and misinterpretations; the most important corrections are acknowledged in the relevant individual entries. jan løve’s name compilations are on file (in danish) on the website of the danish arctic institute (www.arktisk institut.dk: østgrønlandske stednavne). niels henriksen (geus and birkerød, denmark) has been a constant source of support and encouragement throughout the compilation process for this volume. he was also leader of all the ggu/geus geo logical mapping expeditions in which i have participated. he has kindly read large sections of this vo lume, and provided many helpful comments and suggestions. many individuals have kindly provided valuable information on place names given by themselves or by others during expeditions in which they participated. i am particularly grateful to: svend bendix-almgren, john cowie, peter r. dawes, henrik elling, j.d. (di dom) friderichsen, john haller, colwyn jones, david malmquist, arne noe-nygaard, n.e. odell, fritz h. schwarzen bach, cordelia stamp, w. stuart watt and anker weidick. the support of velux fonden is gratefully ac know ledged, and is provided under the statutes of the fund that encourage and give financial support to active pensioneers. finally, i would like to acknowledge the very helpful suggestions of the two reviewers, christopher ries and ian stone. 12 professor n.e. nørlund, director of the geodætisk institut (geodetic institute), wished to solve the problems of names given in various languages by expeditions of different nationalities, and also the use of east and west greenlandic dialects, as east greenland orthography diverges from that of west greenland. nørlund therefore took the initiative to form the place name committee for greenland (sted navne kommis sionen or stednavneudvalget), under the auspices of the scientific commission for greenland, with the aim of ratifying place names in greenland. the initiative was prompted by the introduction of the regulations of journeys to and from greenland issued by the ministry for shipping and fisheries on 7 august 1930, of which section viii states (in translation): expeditions that wish to bestow place names on localities visited, must send proposals to the danish government, who will make the final decision. the first meeting of the committee was held on 6 june 1933. the members included prominent green land administrators and scientists: jens daugaardjensen, lauge koch, niels erik nørlund, knud ras mussen, william c. thalbitzer, f.o. jørgensen and hother ostermann. a second meeting was held on 15 november 1933. officially the place name com mittee for greenland was established on 1 fe b ru ary 1934, when the danish state department issued a regulation announcing the establishment of a place name committee, and stated that no place names given to greenland localities by expeditions would be recognised by the danish state until they had been approved by the committee. four meetings of the place name committee were held in 1934, eight in 1935, and regular meetings were held subsequently until interrupted by the war years. one of the early decisions was to establish a subcommittee with the task of considering for approval all place names that had hitherto been used on pub lished maps. in respect of east greenland, the systematic listing of published place names, and their approval or deletion continued until the 1940s. in the post-war pe riod, up to 31 december 1983, the sub-committee con tinued to approve, modify or reject newly proposed place names, of which final approval was then made by the full committee. the minutes of the place name com mittee for greenland from 1933 onwards were for m erly accessible at the danish geo detic institute, and the main activities and conclusions of the com mittee and its sub-committees rele vant to northern east greenland are summarised here. greenland was granted home rule in 1979, and took over responsibility for its place names on 1 jan uary 1984; the place name committee archives are now in nuuk, green land. one of the early difficulties facing the committee was the significant differences between the west green land dialect and that of east greenland, and the consequent varied spelling of greenlandic place names. at the first meeting of the committee in june 1933, professor william c. thalbitzer, the acknowledged expert on the east greenlandic dialect, argued strongly for preservation of the east greenland forms, rather than the ‘incorrect’ variations introduced in the ammassalik / tasiilaq region by west green land in ter preters such as hansêrak. at the third meeting it was commented that up to six dialect variations might be required to accurately reflect local usage. jens daugaard-jensen, director of grønlands styrelse (the greenland administration), expressed his preference for the general application of samuel kleinschmidt’s orthography as practised in west greenland, a view supported by c. wilhelm schultz-lorentzen who pro phesied (incorrectly) that there would be a general movement towards a common (west greenland) dia lect throughout greenland. thalbitzer threatened to resign from the committee at the 14th meeting in november 1935, partly on the grounds that his views on preservation of dialect forms were repeatedly overruled by other committee members, and partly due to disagreement on the principles for approving future place names. he confirmed his resignation at the 17th meeting in february 1937. the east green land dialect continues to thrive today in the east greenland towns of illoqqortoormiut (ittoq qor toor miit) and tasiilaq / ammassalik and outlying settlements. the preferred east greenlandic spelling of the name of the town scoresbysund is ittoqqortoormiit, and this is the spelling used by the inhabitants and on the official website (www.eastgreenland.com), but it is the west greenland spelling illoqqortoormiut that appears on official maps of greenland. early meetings of the committee were marked by at times acrimonious discussion on the commemora13 official place names in greenland tion of living persons in place names. the third meeting agreed that commemoration of living persons should be reduced to a minimum. an analysis of pre vious practice in east greenland presented at the sixth meeting revealed that 69% of the names pro posed by a.g. nathorst in 1899 commemorated persons, and that 54% of his names had been given after persons then still alive. the corresponding figures for g.c. amdrup’s 1900 expedition to southern east greenland were 80% and 74% respectively. at the fourth meeting reference was made to a somewhat caustic letter by ejnar mikkelsen, who had drawn attention to some of the names on the 1932 edi tion of the 1:1 million scale topographic map com piled by lauge koch (geodætisk institut 1932) that commemorated persons without the remotest con nection with greenland (e.g. anna sten gletscher and gerda gletscher that were named after actresses; see also fig. 15). the large number of names arising from the activities of 1931–34 treårsekspeditionen had also attracted unfavourable press comment because so many had been given after living persons. most of the names applied to geographical fea tures during the 1931–1934 treårsekspeditionen by lauge koch, and the scientists working under his lead ership, were published between the regulations of 1930 and 1934, and thus essentially prior to establishment of the place name committee. lauge koch there fore argued, at the ninth meeting of the committee in february 1935, that the decree of 7 august 1930 was the authority. since this stated that names should be placed before the government for approval, and since koch was the appointed police authority in east greenland during treårsekspeditionen, then he was also (in his view) to be considered the government authority and thus could approve his own names. as an additional argument for the blanket approval of all names given during treårsekspeditionen, koch cited the usage of his maps as documentary evidence at the international court of justice in the hague, during the danish–norwegian controversy over the sovereignty of east greenland. at the eighth meeting of the committee in feb ruary 1935, discussion on the commemoration of living persons in place names concluded with the recommendation that they should be avoided as far as possible, although this might prove difficult in practice. it was proposed that the committee should de cide in individual cases, by vote if necessary. this decision was soon brought into effect with, at the 10th committee meeting in may 1935, the rejection of many names proposed by j.g. jennov (director of the danish trapping company nanok) and ejnar mik kelsen. in rejecting jennov’s names it was incorrectly stated that they had been given after the act of 1 feb ruary 1934; jennov argued that many of his proposed names were given during the 1932 gefion expedition and were in common use amongst danish hunters. jennov’s names were rejected for the third time in 1940, when a sub-committee suggested alternatives for three of jennov’s disputed names (tuxen ø, engel hardt sund and frieda sø), which became nanok ø, jægersund and gunner andersen sø. nume rous sub se quent attempts, by various expeditions, to introduce names obviously given after living persons were rejected. however, the regulations were often circumvented, for example by geologists of lauge koch’s expeditions who would include personal names on their names lists with the discrete explanation ‘girl’s name’. occasionally the place name committee ap pears to have simply turned a blind eye to such pro posals, and for example approved the names ebbe sø, eigil sø and winston bjerg proposed by the 1952–54 british north greenland expedition, although they obviously commemorated ebbe munck, eigil knuth and sir winston churchill (all then alive). exceptions to the ‘living person’ rule are only officially allowed for the danish royal family, a practice that has con tinued to the present day: e.g. dronning margrethe ii land (1990, on the occasion of the danish queen’s 50th birthday), qeqertaq prins henrik (2004, on the occasion of the 70th birthday of the prince consort – the danish queen’s husband), kronprins frederik land (2008, to commemorate the military service of crown-prince frederik in the sirius sledge patrol). at the sixth meeting of the committee in march 1934 it was agreed to establish a sub-committee, the tasks of which were to go through all published da nish and foreign maps, and to make decisions on the danicised form of names to be approved for official usage. name lists were drawn up for consideration by the geodetic institute, divided up for convenience into degrees of latitude, and with the names numb ered consecutively. this system was also to apply to future proposed names, with the number given to each name following it throughout the entire ap proval process. the first meeting of the sub-committee was in january 1935. some principal decisions had already been made by the full committee, such as the usage of west greenlandic spellings for localities in east green land (as noted above), and the usage of the letter q for 14 the special greenlandic ĸ introduced by samuel kleinschmidt. amongst other proposals, usage of the danish aa form was preferred to the swedish å (a decision reversed in 1948). hyphens were to be avoided, such that composite names such as zoolog-dalen were to be given in one word as zoologdalen. names given after persons were to be expressed in two or more words (e.g. milne land, not milneland. in practice it was the sub-committee that made recommendations on place names to be ap proved or rejected, their proposals then being placed before the full committee. at the first meeting of the sub-committee the new names appearing on the maps produced by norges svalbardog ishavsundersøkelser (nsiu) on scales of 1:200 000 and 1:1 million in 1932 were considered, and with only a few exceptions all were rejected, on the grounds that they were politically motivated. a similar fate was to be meted out to the 299 new names given on the norwegian 1:100 000 scale maps covering clavering ø, geographical society ø and jordan hill (lacmann 1937). although the committee ad mitted that lacmann’s maps contained significantly more detail than the best existing danish maps, the procedure for approval of new names by danish authorities had ‘not been followed’. in the event, a few names used on clavering ø were allowed by the subcommittee in 1939, but all others were rejected. the sub-committee approved long lists of names given after localities in denmark, notably those pro posed by the surveyors of the geodetic institute. some lists of names proposed by swiss geologists or british scientists were adjudged too foreign-sounding, even in danicised versions, and were rejected en tirely or replaced by the sub-committee’s own sug gestions. other lists of equally foreign-sounding names were approved. although the committee as early as 1937 had expressed the view that large num bers of foreign-sounding names were to be avoided as far as possible, the rule was inconsistently applied. the indiscriminant usage of the genitive ‘s’ in place names was raised at the eighth meeting of the committee in february 1935. following the recommendations of the sub-committee, already published names that did not use the genitive ‘s’ were considered to have won recognition in that form, whereas newly proposed names should use the genitive form except where circumstances argued against it. however, in practice usage continued to be inconsistent, and the problem was raised again in the 1960s and 1970s when it was realised that charts published by the danish hydrographic office used one form, and the geodetic institute map sheets the other form. one of the last decisions of the place name committee, to restore consistency, was to remove all the genitive ‘s’ endings previously approved. to resolve the general problem in greenland of the use of both danish and greenlandic names for the same feature, particularly as applied to towns and settlements, the principle of officially approving both danish and greenlandic place names was established. when greenland acquired home rule in 1979 there was a subtle change in the ‘double’ name giving, with the greenlandic town names taking precedence over the danish equivalent. however, while in some cases the danish town names gradually fell into disuse, in other cases the danish town names were ‘officially’ abandoned by decision of the local town council leaving just the greenlandic names. in some towns this policy went to the extremes of deliberately replacing all former danish street names with greenlandic alternatives. however, maps that show the original double names for greenland towns still appear in the most recent greenland atlases (berthelsen et al. 1989; jakobsen et al. 2000). the regulations concerning travel to and from greenland were revised in 1939 and 1948, and a regulation of 11 april 1949 re-organised the place name committee, with eske bruun (then head of the green land administration) as chairman. the responsibilities of the new committee were essentially identical to those of the original committee. one of the early initiatives of the new committee was to undertake the systematic collection of greenlandic place names used by local populations in greenland. this process began in 1949, and in 1955 a two-man party from the geodetic institute visited scoresbysund / illoqqor toor miut (ittoqqortoormiit) and collected a total of about 190 names used locally, the majority being of the characteristic descriptive type. following revision of the danish orthography in 1948, the changes proposed were also applied to danicised place names in greenland. the main change was that of the da nish aa to å, although geodetic institute map sheets continued the old usage until 1954. in 1973 a major revision of the greenlandic ortho graphy was implemented (greenland spelling re form), with the new system notably abandoning all the accents on letters (introduced by samuel klein schmidt as an aid to pronunciation). the modern written language, and spelling of place names, makes extensive use of double vowels and consonants. since 15 the existing geodetic institute map sheets, and other atlases, used the old-style spelling, a systematic database of all place names in greenland was compiled by the geodetic institute in 1986–1987, that includes both the old and new greenlandic spellings for place names in greenland. both old and new spellings of greenlandic approved names in northern east green land are given in this volume, with the main entry under the new spelling, with cross-references for all the old spellings. in 1979 greenland was granted home rule, which meant that greenland acquired a special status within the kingdom of denmark, with its own parliament in nuuk in west greenland. as a result of this major change many responsibilities previously carried out by denmark on behalf of greenland were transferred to greenland. on 31 december 1983 the place name committee for greenland was disbanded, and responsibility for place names in greenland was transferred to grønlands sprognævn, today the nunat aqqinik aalajangiisartut / grønlands stednavnenævn / green land place names committee. in 2009 greenland was officially granted selfgovernment (selvstyre), a further measure of independence from denmark, with exceptions in respect of foreign policy and defence, but still with a substantial annual subsidy of 3400 million danish kroner (c. $637 million). 16 this section comprises brief summaries of all activities in northern east greenland from c. 2500 bc up to the present day. these activities range from large scientific expeditions with more than 100 participants to minor tourist visits by a few persons. due to the remote and isolated situation of northern east greenland, virtually all visiting groups need to be self-supporting and are therefore characterised as ‘expeditions’. the various activities are presented chronologically, with brief information on the nature of the objectives and results of scientific investigations, and with particular emphasis on place names proposed by the participants. in general the name of the expedition is given in the original language, followed by the expedition name in english (where relevant) and the name of the leader. in 1979, greenland was granted home rule, and took over many of the responsibilities formerly car ried out by denmark on its behalf. from 1989 until 2009 the danish polar center (dpc) undertook the issue of permits to visit northern east greenland and the north-east greenland national park. one of the conditions of the permits was that a report should be submitted to dpc, but some expeditions have failed to submit reports and the expedition list that follows is therefore incomplete. many of the modern activities, from about 1961 onwards, are recorded only in unpublished expedition reports deposited with the organisations that supported the activity, or from 1989 with the danish polar center. where such reports have been located at the british mountainering council in manchester [bmc report archive], royal geographical society in london [rgs report archive] or the danish polar center in copenhagen [dpc report archive], this is indicated at the end of the activity description. in 2009 greenland acquired a further degree of independence from denmark, and from 2010 permission to visit northern east greenland must be re quested from the ministry of domestic affairs, nature and environment of the government of green land in nuuk. note that in the following accounts of activities officially authorised place names are in normal type, whereas unofficial place names, or unapproved variations of names, are given in italics. the names of ships are given in capitals, e.g. the hopewell. note that hunting trips made by the residents of scoresbysund / illoqqortoormiut (ittoqqortoormiit) are not included in this volume, although such tours may extend northwards into the north-east green land national park and southwards along the blosseville kyst. local excursions organised by the travel agents nanu travel aps for groups visiting scoresbysund / illoqqortoormiut (ittoqqortoormiit) are in general outside the scope of this volume. similarly, excursions by personnel from danmarks havn weather station (or the former weather and/or radio stations at daneborg and kap tobin) are not generally on public record. the numerous scientists that visit the zackenberg ecological research opera tion (zero) are mainly involved in projects in the vicinity of the research station, but a few projects range more widely afield (see meltofte et al. 2008); a few zero projects merit mention below, but most projects are included in the general descriptions of activities – see: ‘1997–present zackenberg ecological research operations (zero)’. cruise ships have occasionally visited east greenland since the early 1970s, and many cruise organisations now include regular visits to east greenland in their schedules; up to 17 ships annually have been recorded carrying a total of about 1000 passengers – see: ‘1998–present: nuna travel aps’. with few exceptions these cruises are not individually listed here. the sirius sledge patrol covers a total of 20 000 km on patrol with dogsledge teams in northern east greenland and north greenland during the winter and spring every year. details of these patrols are confidential, but p.s. mikkelsen (1986, 2005) has provided an informative and well-illustrated account of his own experiences with the sirius sledge patrol. 17 exploration history of northern east greenland pioneer exploration and discovery: c. 2500 bc – ad 1912 c. 2500 bc – c. 1823 inuit (palaeoeskimo) immigrations about 4500 years ago, long before european whalers and explorers set foot on the east coast of greenland, the entire region had been settled by inuit (palaeo eskimos). the independence i culture, which is closely related to the saqqaq culture of west greenland and the early pre-dorset culture of canada, had spread from ellesmere island (canada) across north greenland and down the coast as far as scoresby sund (70°n). the independence i people remained in east greenland for up to 600 years (bennike et al. 2008). about 1100 years later a new wave of inuit (pa laeo eskimos), the greenlandic dorset, retraced their pre decessors’ footsteps. both cultures depended for their existence on musk oxen, seals, hares, birds and fish. the tent rings of the greenlandic dorset are widely distributed along the coast of east greenland, with a concentration in dove bugt and on île de france (now qeqertak prins henrik). the green landic dorset people lived in east greenland from about 800 bc to 0. about ad 1200 the ancestors of the present day greenlanders, the thule culture, reached greenland, and via north greenland soon populated the entire coast of east greenland. they were whale-hunters and possessed skin boats (kayaks and umiaks), but also depended on musk oxen, seals, hares, birds and fish (larsen 1970). the last remnants of this population north of 69°n latitude may have been the group of 12 encountered by douglas clavering at clavering ø (74°15´n) in 1823 (clavering 1830). ruins of their winter houses are common throughout east green land. c. 1000–1250 norse (viking) voyages the icelandic sagas include accounts of a number of voyages to greenland, although most of the place names recorded have usually been identified with locations in south or west greenland (rafn 1845). some names have appeared in a variety of positions on old charts which were based partly on interpretations of the sagas (egede 1818; steenstrup 1886, 1889; bjørnbo 1911; trap 1928; see frontispiece). how ever, tornøe (1935, 1944) has argued that places described in landnámabók, eirik raudes saga, tor finn karls evnes saga and other sources might have been situated in east greenland. north of latitude 69°n tornøe suggests locations for bláserkr, breidi fjórdr, finns búdir, greipar, krosseyjar and òllumlengri. apart from blá serkr (now rigny bjerg) their positions are debateable, and none of them have acquired the status of approved names. òllumlengri or ollum lengri fiordr is said to have been discovered by norse voyagers from iceland in 1194 or 1195, and here they for many years hunted seals, walruses, narwhales and bears. gustav holm (1925, 1926) considered their description of the ‘fjord longer than all other fjords’ admirably fitted the present day scoresby sund, a viewpoint supported by tornøe (1944) and ejnar mikkelsen (1989). scoresby sund with its inner branch of nordvestfjord is in fact the longest fjord in the world. the icelandic annals also refer to the discovery in 1194 of svalbardr, or svalbarda í hafsbotn, the ‘country of the cold coasts’, which some authorities identify with the scoresby sund region (70°–72°n) of east greenland (rafn 1845; ryder 1892; holm 1926), others with jan mayen (wordie 1922) or spitsbergen (tornøe 1935, 1944). svalbard is today the official name of the group of islands including spitsbergen that were placed under the sovereignty of norway by the treaty of paris in 1920. direct evidence of norse visits to east greenland north of latitude 69°n is limited to finds in inuit graves at scoresbysund of silver buttons and beads (storgaard 1926) and of an ornamented bone comb (thalbitzer 1909); these have been argued by tornøe (1944) to indicate some contacts between the norse inhabitants of iceland and the former inuit population. 1607 henry hudson’s voyage in 1607, henry hudson was sent out by the muscovy company with a crew of 11 on the hopewell to seek a passage to japan and china across the north pole. he sighted the coast of east greenland on several occasions between latitudes 68° and 74°n, and on 22 june 1607 lay off hold with hope (73°30´n). the only account of his observations is reproduced in asher (1860) and purchas (1906) – “it was a mayne high land, nothing at all covered with snow: and the north part of that mayne high land was very high mountaynes .... wee thought good to name it, hold with hope, lying in 73. degrees of latitude” (asher 1860, p. 3; purchas 1906, p. 297–298). 18 hold with hope is the oldest place name currently in use in northern east greenland. while hudson failed in the main purpose of his voyage, his accounts of the abundant whales in the waters near spitsbergen are said by many authors to have led to the development of the northern whale-fishery; other writers give the credit to nicholas woodcock’s 1612 voyage (see below). c. 1614 – c. 1910 northern whale-fishery until the pioneer charting of the coast of east green land by william scoresby jr. in 1822, the only information on the region north of latitude 69°n came from the chance sightings of whalers. british whalers began to sail to spitsbergen waters after nicholas woodcock’s successful voyage in 1612, and as a result of their success were soon joined by dutch whalers, and subsequently by french, spanish, danish and others. whales became scarce in the bays of spits bergen after 1630, leading to a temporary decline in british whaling. after 1720 whales were then sought along the edge of the east greenland pack ice. revival of british whaling about 1750 was linked to the introduction of a government bounty. fluctuations in whal ing returns, especially in the british trade, were in fluenced by variations in the bounty (which lasted until 1824), the attacks of hostile privateers, the weather conditions and whale migrations. in view of the numbers of whalers engaged in the fishery, there were probably numerous sightings of the greenland coast, but records are few. no deliberate attempts were apparently made to penetrate the ice belt before 1822, the general opinion among whalers up to about 1818 being that the land was inaccessible (scoresby 1823). a note on an italian map from 1690 by coronelli records that the dutch sighted the coast of east greenland at about 79°n in 1614, and that broer ruys reached land and observed gael hamkes land at c. 73°n in 1654 (bobé 1928). a collection of dutch charts, ‘de groote nieuwe zee-atlas door gerrit van keu len’ from 1706, includes a chart recording the discovery of ‘t’land v. broer ruys’ in 1655 at 73°30´n, ‘t’bay v. gale hamkes’ in 1654 at 74°n, ‘t’land v. adam’ in 1655 at 77°n and ‘t’land v. lambert’ in 1670 at 78°30´n. nearly all these names were preserved by subsequent explorers, and were later approved in danicised form. in 1761, a danish whaler, volquaart boon, aboard a dutch or german ship, followed the east greenland coast from 76°30´ to 68°40´n, and at about latitude 70°20´n was dragged by a strong current into a wide and deep fjord, the present scoresby sund (bobé 1928). other whalers known to have sighted the coast, usually reported as gale hamkes land, include die frau maria elisabeth in 1769, de sankt peter in 1773 and willemina in 1777 (ryder 1892). in 1798, british cruisers had captured the dutch whaling fleet, and by the early 1800s the northern whale fishery was largely in british hands. a series of prosperous whaling years lasted until about 1826, although with a progressive shift in interest from the greenland sea to the davis strait (offshore west greenland). william scoresby sr. and his son had notable success in east greenland waters, and their search for the declining whales led to attempts to penetrate the pack ice. william scoresby jr. sighted land at 74°n in 1817, and in 1821 observed the coast from 74°30´ to 73°30´n (manby 1822); william sco res by sr. also followed the coast in 1821 from 74° to 70°n (scoresby 1823). however, all these observations were from a great distance, and it was only in 1822 that william scoresby jr. came close enough to the coast to construct a chart (see below). other whale fishers also approached the coast and good catches were often made. from about the 1750s whalers had begun to take seals in increasing numbers, hamburg and altona ships taking 50 000–60 000 in the greenland sea in 1787. as whaling declined, sealing gained in importance, scottish ships beginning intensive sealing in 1831, were joined in 1847 by norwegian sealers who subsequently dominated the trade (see below). whaling in east greenland waters was maintained largely due to the enterprise of a few notable whaling skippers. following the retirement of the scoresbys’ after 1822, the gray family of peterhead were most celebrated, with their equally notable ships, active, eclipse and hope. they were amongst the few to make paying voyages to the greenland sea in the 1870s, and the peterhead fishery ceased with the retirement of david gray in 1891. tom robertson was among the last to seek whales off east greenland, and made regular voyages from 1895 until 1907 with the active and balaena with moderate success, and occasionally reached land. in 1899 he assisted a.g. nathorst’s expedition, and took home 10 musk oxen. the effective end of the greenland whale fishery is placed at about 1910 (lubbock 1937; jackson 1978). 19 1822 william scoresby jr.’s whaling voyage william scoresby jr. and his father were important figures in the history of arctic whaling, but were also natural scientists, and even while engaged in the search for whales concerned themselves with scientific observations of all kinds (stamp & stamp 1975). william scoresby jr. became one of the leading au tho rities on magnetism, especially on marine com passes and their deviation, published many articles on a variety of subjects, and has been considered a founder of arctic science and the beginnings of ocean ography. one major result of scoresby jr.’s whaling career was his celebrated two volume ‘an account of the arctic regions’ (scoresby 1820), and the journal of his 1822 voyage which brought back for the first time anything approaching accurate information on the fjord region of east greenland (scoresby 1823). between june and august 1822, william scoresby jr. on the baffin was close to land on numerous occasions, sometimes in company with his father on the fame, sometimes with other whalers – up to 20 or 30 whalers were at times reported in sight. william scoresby jr. succeeded in laying down a chart of the east greenland coast between latitudes 69° and 75°n, the original of which is now in whitby museum (eng land). the most accurate portion is that from 70° to 72°30´n, where landings were made at kap lister, neill klinter, kap brewster and kap moorsom (fig. 3), the first landings recorded by european visitors. areas farther north were observed from a distance. scoresby (1823) recorded geological, botanical and zoological observations. scoresby sund was given its name after william scoresby sr., described as the first to enter the sound, and hurry inlet was explored. one of the most important results of scoresby’s survey was a correction of the serious errors of longitudes, placed 7° to 14° too far to the east on earlier charts. subse quent explorers have had little difficulty in recognising the features scoresby laid down, and nearly all of scoresby’s 80 place names have survived. however, a few earlier dutch names were misplaced by scoresby, and some of his capes subsequently proved to be mountains standing well back from the coast (white 1927). the majority of scoresby’s place names were given after his friends, notably including a number of scientists from edin burgh who had encouraged his scientific interests. 1823 voyage of douglas clavering and edward sabine on the griper the british board of longitude decided that edward sabine’s pendulum observations should be continued to the most northerly latitude possible, and appointed douglas clavering as captain of the griper for a voyage to spitsbergen and greenland in 1823. ed ward sabine’s pendulum experiments were aimed at determination of the earth’s magnetic field and the shape of the earth, and for this purpose he had travelled widely in america and africa. after completion of observations in spitsbergen in 1823, course was set for greenland. an attempt to penetrate the ice belt at 77°n failed, and the coast was eventually reached at about 74°n. an observatory was set up on what was subsequently called sabine ø (74°35´n) on 13 august, and the pendulum experiments successfully completed (sabine 1825). meanwhile a boat journey was made by douglas clavering to the present clavering ø 20 fig. 3. part of the chart of the east greenland coast drawn up by william scoresby jr. in 1822, showing the numerous features that he named in liverpool land (liverpool coast) and adjacent areas. from: scoresby (1823). (74°15´n), where the only recorded meeting with the last remnants of the thule-culture inuit was made on 16–19 and 23–24 august (clavering 1830; ryder 1892). clavering also explored and named loch fyne (73°45´n). in the course of the voyage clavering, with his midshipman henry foster, surveyed the coast be tween 72°30´ and 74°n, joining up with the 1822 ob servations of william scoresby jr. all of clave ring’s 18 names have survived. most were given for scottish localities and friends, while the islands on which the pendulum experiments were carried out are com memorated as the pendulum øer. 1831 albert haake and the bremen albert haake, sailing on the bremen, is reported to have made a landing in east greenland at about 74°n in july 1831, and reported a broad strip of ice-free water along the coast (ryder 1892). 1833 jules de blosseville and la lilloise jules de blosseville was a french naval officer who in 1833 had command of the brig la lilloise, and the task of maintaining order among the whalers and fishing vessels around iceland. on 29 july he sighted the coast of east greenland between 68° and 69°30´n that now bears his name. he returned to iceland to dispatch a report and sketch-map of his discoveries, and on 5 august set sail again to continue his observations, but vanished without trace with his crew of 80. his map included a number of names mostly given for ministerial officials (fig. 4), and while exact identification of his named features was often not possible, georg carl amdrup’s 1898–1900 expedition pre served many of them (blosseville 1834; amdrup 1902a, 1902b). only four were given for features north of latitude 69°n, of which two names survive on modern maps – rigny bjerg and d’aunay bugt (c. 69°n). 1847–1959 norwegian fishing and hunting voyages norwegian sealers made their first appearance in the greenland sea in 1847, and within a few years at tained a dominance of the trade. sealing reached its height in the 1850s, and in one season 40 ships took 400 000 seals. norwegian landings on the coast of east greenland can be dated back to 1889, a poor sea ling season, when the hekla captained by ragnvald knudsen visited the coast between 73°30´ and 75°30´n. the hekla returned home with a substantial catch of more than 2700 seals, 267 walrusses, 9 bears and 24 musk oxen (knudsen 1890; solberg 1929). in subse21 fig. 4. this map shows jules de blosseville’s observations in 1833 north and south of latitude 69°n, on what came to be known as blosseville kyst after he was lost in the ice with the crew of la lilloise. this was the map sent home to france from iceland, after which blosseville returned to the east greenland coast with the intention of continuing his observations (blosseville 1834). very few of blosseville’s place names have been preserved, except for rigny bjerg (mt. rigny) that later proved to lie just north of latitude 69°n. 28˚ 27˚ 69˚ quent years norwegian sealers periodically followed the hekla’s example, visiting the coastal waters to supplement their catch of seals. isachsen & isachsen (1932) record 142 visits by nor we gian ships between 1889 and 1931, numbering usually one to four each year, but with eight in 1900 (isachsen 1922). catches were sometimes notably large, that of the aspø in 1898 including 66 bears, those of the søstrene and the spidsbergen in 1899 including 79 and 69 musk oxen respectively, and the first two live musk-ox calves, while the spidsbergen in 1901 took 46 wal russes. in 1908–09 the first nor wegian overwintering expedition was led by severin liavaag, followed in 1909–10 by vebjørn landmark’s expedition. nor we gian ships made something of a speciality of bringing live musk oxen to europe for sale to zoos, and alendal (1980) records that 290 musk oxen were brought back between 1899 and 1969. after the signing of the danish–norwegian treaty on east greenland in 1924 (østgrønlands traktaten, see below), a succession of norwegian and danish fox-trapping expeditions wintered in east greenland, some state-supported while many others were private initiatives. most are briefly described individually be low. a general account of norwegian hunting up to 1939 is given by rogne (1981), and a detailed account of all danish and norwegian trapping activities by p.s. mikkelsen (1994, 2008). in 1946, after world war ii, norwegian hunting was resumed under the auspices of arktisk nærings drift and hermann andresen (see also below). by 1959 hunting had virtually ceased following withdrawal of state subsidies and falling skin prices. as a consequence of reduced norwegian activity, and other factors, denmark availed itself of the termination clause in the danish–norwegian treaty, which expired on 9 july 1967. 1869–70 die zweite deutsche nordpolarfahrt (the second german north pole expedition): karl koldewey this expedition was organised on the initiative of the noted german geographer august petermann, who 22 fig. 5. the discovery of petermann bjerg (2970 m) by karl koldewey’s 1869–70 expedition. the mountain is on the skyline at the centre of the drawing, and was observed by julius payer after climbing sonklargletscher to a vantage point on the icecap north-east of payer tinde (the peak at left). the fjord visible is the inner part of kejser franz joseph fjord, first completely explored by a.g. nathorst’s 1899 expedition. from: verein für die deutsche nordpolarfahrt in bremen (1873–74). had suggested an attempt be made to reach the north pole along the coast of greenland or spitsbergen. a reconnaissance expedition led by karl koldewey on the grønland was sent out in 1868, failed to penetrate the pack ice off east greenland, but eventually reached spitsbergen. based on this experience a larger-scale expedition was organised, and in june 1869 the steamer germania, especially built for the voy age, together with the schooner hansa, set out for east greenland. the germania reached land at 74°n. however, the hansa was crushed in the pack ice and sank off the coast of liverpool land (71°n), the crew drifting on an ice floe down the coast, rounding kap farvel (59°46´n) and eventually reaching land near the settlements in west greenland. the germania was captained by karl koldewey, and the ship’s officers included the austro-hungarian lieutenant julius von payer, ralph copeland as sur vey or, carl börgen as meteorologist and adolph pansch as surgeon. after failed attempts to penetrate northwards along the coast with the ship, the ger mania anchored in germania havn on sabine ø (74°32´n) where it overwintered. in the autumn of 1869, sledge journeys were made to fligely fjord, kuhn ø, clavering ø and tyrolerfjord (74°–75°n). in the spring of 1870, two sledges and 10 men were sent northwards along the unknown coast and reached germania land at 77°n. further sledge journeys were made to ardencaple fjord, shannon and clavering ø. in the summer of 1870, attempts were made to press northwards with the germania, but without success, and the expedition turned southwards to dis cover and partially explore kejser franz joseph fjord (73°15´n). the local ice cap adjacent to payer tinde was climbed, from the top of which petermann bjerg was sighted far inland to the west (fig. 5). although the expedition failed to reach the north pole or to demonstrate a practical route, it made important geographical discoveries and mapped large parts of the coastal region of east greenland between 73° and 77°n. important meteorological, geological, botanical and zoological observations were made. this ex pedition was the first to report musk ox in east greenland. the detailed maps of the expedition record about 125 new place names (fig. 6; verein für die deutsche nordpolarfahrt in bremen 1873–74; koldewey 1874; payer 1876, 1877), nearly all of which survive on mo dern maps. the names proposed were evidently the work of a committee and incorporate many suggestions of august petermann (see e.g. verein für die deutsche nordpolarfahrt in bremen 1870–76). most were given for prominent german scientists, the officers and scientists of the ships, and colleagues who had assisted or promoted the expedition. others were given during the expedition and commemorate incidents (e.g. stormbugt), or the ap pearance of features (e.g. eiger, tyrolerfjord, teufel kap). 23 fig. 6. segment of the map just north of 74°30′n produced by karl koldewey’s 1869–70 expedition. from: verein für die deutsche nordpolarfahrt in bremen (1873–74). the expedition wintered at germania havn (germania hafen) on the south side of sabine ø (sabine i.). 1879 orlogskonnerten ingolf ekspedition i dan marksstrædet (the ingolf expedition to danmark strait) the danish schooner ingolf captained by a. mourier was dispatched in 1879 to undertake hydrographical observations in danmark strait. it came suf fic iently close to the east greenland coast to sketch many features between 65° and 69°n (mourier 1880). only few are relevant to this account, and in clude a more accurate placing of jules de blosse ville’s mont rigny (rigny bjerg). 1891–92 den østgrønlandske expedition (the east greenland expedition): carl ryder lieutenant carl ryder was appointed leader of an 11man danish government-sponsored expedition to east greenland, which sailed from copenhagen in early june 1891 aboard the norwegian sealer hekla, captained by ragnvald knudsen. a direct route through the ice pack to scoresby sund proved im prac tical, and a detour was made to the north, the coast being reached in the vicinity of hold with hope (73°40´n) on 20 july, and the mouth of scoresby sund (70°20´n) on 31 july. after entering scoresby sund, a visit was made to kap stewart, the site originally planned for the wintering station, but this proved not to be suitable. from a vantage point on neill klinter it was observed that hurry inlet was not a channel as depicted by william scoresby jr. in 1822, but a closed fjord. sail ing westwards into the unknown inner reaches of scoresby sund, a small enclosed harbour (hekla havn) was discovered on danmark ø, and became the winter harbour for the expedition and ship. from hekla havn journeys were made by motor boat into gåsefjord, føhnfjord, rødefjord and nord vest fjord, the first explorations by europeans, as well as along the coast of jameson land. in spring 1892, several sledge journeys were made. the first revisited føhnfjord and rødefjord, and discovered rypefjord and harefjord. the second pene trated to the inner parts of vestfjord. subse quently journeys were also made to sydbræ and the inner parts of gåsefjord. details of the journeys are found in the official report of ryder (1895), the diaries of ragnvald knudsen published in edited form by giæver (1937), and the diaries of lieutenant helge vedel (gulløv 1991). in august 1892, the hekla left hekla havn, with a stop being made at kap stewart where a depot house (ryders depot) was constructed. the hekla then sailed via iceland to ammassalik, and after a short visit returned to copenhagen. in addition to exploration and mapping of the inner ramifications of the scoresby sund fjord system, significant botanical, zoological and geological observations were made (fig. 7). about 50 new place names are recorded, nearly all of which were given for natural features, incidents and the animal life of the region. 24 fig. 7. part of the geological map of the inner scoresby sund region produced during carl ryder’s 1891–92 expedition. the expedition wintered at hekla havn on the south side of danmark ø (danmarks ö) from: bay (1896). 70˚ n 25˚ w 1898–1900 carlsbergfondets expedition til østgrønland (the carlsberg foundation ex pedi tion to east greenland – often called the 1898– 1900 amdrup expedition): georg carl amdrup this was a three-year danish expedition, but the work of the first two years (1898–1899) was entirely in the ammassalik region (65°–66°n), and it was only in 1900 that it turned its attention to surveying and exploration of the almost unknown coast extending northwards between ammassalik and scoresby sund. the antarctic left copenhagen in mid-june 1900 with an 11-man expedition led by g.c. amdrup that reached the coast of east greenland at lille pendulum (74°40´n). turning southwards the expedition reached kap dalton (69°25´n) on 18 july and there divided into two parties (amdrup 1902a). after building a depot house just to the north of kap dalton, amdrup set off southwards with a crew of three in an 18-foot open boat along the virtually unknown blosseville kyst. ice conditions were more favourable than expected, and the expedition suc ceeded in making a rough chart of the coast between kap dalton (69°25´n) and agga ø (67°22´n). ammas salik was reached on 2 september (jacobsen 1900; am drup 1902b). meanwhile, the antarctic with the remainder of the expedition under the leadership of nikolaj hartz explored the islands and fjords north of kap dalton, finding hot springs, and running aground in turner sund (hartz 1902). entering scoresby sund, the ant arctic sailed to the head of hurry inlet where zoo logical and geological excursions were made inland, and carlsberg fjord was discovered. kap brew ster was visited before the antarctic sailed north along the outer coast of liverpool land making several landings and charting further new fjords and valleys. entering kong oscar fjord (72°10´n) an excursion was made into the inner part of forsblad fjord map ped the previous year by a.g. nathorst (see above). the ship then left the coast for iceland, before returning to ammassalik to fetch amdrup’s party. about 30 new names were given for features north of latitude 69°n. some of these commemorate earlier explorers to the coast and danish scientists, while others were given for geological or other characteristics of the localities. j.p. koch (1902), who was responsible for the surveying from the ship, noted that he used all previous names that he could identify with certainty, except for those he considered mis leading. thus, eight of scoresby’s capes were omitted as they appeared to be mountains; some of these names were later transferred to mountains following mapping by james wordie’s expeditions (white 1927). 1899 swedish east greenland expedition: alfred gabriel nathorst a.g. nathorst led two arctic expeditions in search of traces of salomon andrée’s lost balloon expedition (nathorst 1900). the first in 1898 was to spitsbergen, and the second in 1899 to east greenland. the 1899 expedition left stockholm in may aboard the antarctic, met difficult ice conditions, and reached land at scoresby sund (70°10´n) where the head of hurry inlet was visited. when ice conditions improved the antarctic sailed north to the mouth of kejser franz joseph fjord (73°10´n), and 25 fig. 8. segment of a.g. nathorst’s map of his discoveries in the kong oscar fjord region, showing ella ø (ellas ö), named after his wife, and maria ø (marias ö) and ruth ø (ruths ö), named after his two daughters. many of the capes shown were named by nathorst for other members of his family. from: nathorst (1900). followed the entire length of the fjord reaching the inner end for the first time and exploring kjerulf fjord. the connection with kong oscar fjord via antartcic sund was discovered, and the network of interconnecting fjords and islands explored. nathorst chose the mapping of these new territories as more important than other scientific investigations. sur veying was largely undertaken by per dusén with the assistance of f. åkerblom. about 94 new names ap peared on the published maps, many of them given for supporters of the expedition, for expedition mem bers, and notably for members of nathorst’s own family (fig. 8). 1900 till spetsbergen och nordöstra grönland (to spitsbergen and north-east greenland): gustav kolthoff gustav kolthoff led a zoological expedition to spits bergen and east greenland aboard the frithjof in 1900 (kolthoff 1901). the expedition reached land at mackenzie bugt (73°25´n) on 31 july, sailed north to the pendulum øer where mail by tradition was deposited on hvalrosø, and then into kejser franz joseph fjord and moskusoksefjord where two muskox calves were captured. a large collection of birds and animals was taken home, including two wolves. only one new place name was used, tärnholmen for a small island in mackenzie bugt. 1901 baldwin-ziegler depot-laying voyage by the belgica to support the possible line of retreat of the amer ican baldwin-ziegler expedition, which was to make an attempt on the north pole from franz joseph land, depots were laid out by the belgica in specially built huts on southern shannon at kap phil lip broke and on bass rock. the ill-fated baldwinziegler expedition was led by evelyn baldwin and generously financed by william ziegler, but achieved practically nothing. the depots were visited and checked by the magdalena in 1905, in connection with the relief of the 1903–05 fiala-ziegler polar expedition. subsequently the huts and the depots they contained were used by norwegian and danish hunters. 26 fig. 9. part of the map of the observations by the duke of orléans between latitudes 79°30′n and 79°n, made in 1905 aboard the belgica. the northernmost landing was on île de france (from 2004 known as qeqertaq prins henrik). from: orléans (1907b). 1905 expédition arctique du duc d’orléans (arctic expedition of the duke of orléans) this expedition aboard the belgica was led by louisphilippe-robert duke of orléans [1869–1926], with adrien victor joseph de gerlache de gomery [1866– 1934] as captain. after visiting the west coast of spitsbergen, the belgica sailed for east greenland, and off the coast near kap bismarck (76°42´n) met the norwegian sealer søstrene which had reached latitude 77°n and reported ice conditions to be the best its captain had known in 30 years. thus encouraged the belgica pressed northwards along the coast, touching land at 77°35´n, and had reached 78°16´n when stopped by unbroken winter ice. landings were made at several places, and a rough chart made of newly discovered land areas between 77° and 78°50´n. geological, botanical, oceanographic and meteorological observations were also made during the voyage (orléans 1907a, b). soundings were made at 74 locations and the relatively shallow belgica bank was discovered and delineated (barr 2010). l.-p.-r. duke of orléans included 28 new names on his charts, given mainly for members of the orlé ans family, for notable french and belgian ex plorers, and for officers of the ship’s company (fig. 9). few ex planations of the names are given. the duke of orlé ans (1907a) notes with regret that some of the names on his original chart were modified at the request of the danish authorities. thus, his original name terre de france was changed to terre de duc d’orléans, the present hertugen af orléans land. the 1906–08 danmark-ekspeditionen (see be low) had received an advance copy of the orléans chart, and in the course of their explorations re map ped the area in considerably more detail. they record the difficulty of correctly locating the features seen and named by the duke of orléans, and while pre serving as many of the original names as possible, admit that some positions may be incorrect. never theless, it is these positions that have survived on modern maps. in 2010 the cruise ship plancius visited île de france (now qeqertaq prins henrik) with queen paola of belgium and nine descendents of the captain of the belgica,  adrien de gerlache de gomery,  on board. the journey was arranged to commemorate that de gomery had reached the island with the belgica in 1905. a similar but unsuccessful attempt had been made on the 100th anniversary of the event. 1906–08 danmark-ekspeditionen til grøn lands nordøstkyst (danmark expedition to north-east greenland): ludvig myliuserichsen this was one of the largest and most ambitious of early danish expeditions, whose aims were to explore and survey the large unknown region north of kap bismarck (76°42´n) and to link up with the explorations of robert e. peary in north greenland. the expedition numbered 28, including scientists, ship’s crew and three greenlanders, and was led by ludvig mylius-erichsen (friis 1909; amdrup 1913). the expedition sailed from copenhagen on 24 june 1906 aboard danmark, met difficult ice conditions, and reached the coast of east greenland at store koldewey (76°30´n) on 13 august. after sailing north along the coast to íle de france (in 2004 renamed qeqer taq prins henrik), danmark turned south again to danmark havn (76°46´n) which was to be come the expedition base for the next two years (the icao – international civil aviation organi sation weather station ‘danmarkshavn’, spelt as one word, was established on the north side of danmark havn in 1948). during the course of the expedition nearly 200 short and long journeys were made by sledge, boat or on foot. many of these were made during exploration of the islands and fjords around dove bugt south of danmark havn. a meteorological station set up west of danmark havn at pustersvig was manned for a long period by peter freuchen. two journeys were made across the glacier storstrømmen, one via sæl søen to dronning louise land, and the second via annekssø to ymer nunatak. two long journeys were also made southwards along the coast to check the depots at bass rock (74°43´n), and also to deposit the traditional mail. four depot-laying journeys were made northwards in the winter of 1906–07 in preparation for the main spring sledge journeys. on 28 march 1907 a start was made from danmark havn with four parties, in all 10 men and 86 dogs. two of the parties turned back from 80°30´n, surveying on the way and reaching the ship again in late april. at nakkehoved (81°42´n) the two other parties, led by ludvig mylius-erichsen and j.p. koch respectively, parted company. koch’s party went northwards along the east coast of peary land as far as kap bridgman (83°29´n), retrieving peary’s record at kap clarence wyckhoff on the way. returning southwards they unexpectedly met myliuserichsen’s party on 27 may, and then retraced their outward steps to reach danmark havn on 23 june 1907. 27 ludvig mylius-erichsen, niels peter høeg-hagen and jørgen brønlund travelled westwards after parting from j.p. koch’s party, to explore independence fjord and danmark fjord, and were forced by open water to spend the following summer on the west shore of danmark fjord (81°30´n), where they and their dogs suffered badly due to poor hunting. they began their return journey in mid-october, but my lius-erichsen and høeg-hagen died (possibly near nioghalvfjerdsfjorden 79°37´n), while brønlund reached the east point of lambert land (79°09´n) before he also died. two relief parties were sent out to look for the missing party, the first in autumn 1907, and the second in march 1908 that found brønlund’s body and diary. the bodies of mylius-erichsen and høeghagen have never been found, and the precise route followed by the retreating party from danmark fjord to lambert land has remained a lasting topic of speculation (e.g. e. mikkelsen 1913; knuth 1958; lund bye 1984). the expedition sailed back to denmark in august 1908. more than 200 names are associated with the activities of danmark-ekspeditionen in northern east greenland, of which 190 have official status. they record incidents during the expedition, geological characteristics, associations with bird and animal life, while some were named after danish localities, danish personalities and the families of the expedition members. 1908–09 floren expedition: severin liavaag a seven-man hunting expedition on the floren was sent out from the sunnmøre district of norway on the initiative of severin liavaag and the ålesund mer chant hans koppernes, and became the first nor wegian hunting expedition to overwinter in east greenland. the floren anchored in germania havn (74°32´n), and two huts were built nearby, at kap wynn and kap borlase warren. in the winter and spring hunting was carried out between kap herschel and germania havn, and in the summer as far north as shannon (75°10´n). two men were drowned, in cluding liavaag, when they fell through the ice in may 1909 during a bear hunt. the only original pub lished account of the expedition is a diary by brandal (1930), which mentions 15 names used by the hun ters. a brief account of subsequent sunnmøre expeditions is given by rogne (1981). 1909 expédition arctique du duc d’orléans (arctic expedition of the duke of orléans) the duke of orléans, aboard the belgica captained by adrien de gerlache de gomery as in 1905, made a voyage to east greenland, spitsbergen and franz josef land in 1909. in east greenland, difficult ice conditions restricted movements to the area between hold with hope and shannon (73°30´–75°30´n), where they met the surviving members of the 1908– 09 floren expedition (orléans 1911; barr 2010). 1909–10 vebjørn landmark’s expedition a six-man norwegian hunting expedition led by vebjørn landmark was sent out in the 7de juni on the initiative of s.th. sverre of kristiania (oslo). a hunting station was built at kap mary (74°10´n), and a smaller house in germania havn (74°32´n). hunting was carried out between clavering ø and the pen dulum øer in the winter and between jackson ø and shannon in the summer. it was this expedition that in 1910 rescued five members of the 1909–12 alabama expedition from bass rock (see below; e. mikkelsen 1913, 1922). 1909–12 alabama-ekspeditionen til grøn lands nordøstkyst (alabama expedition to north-east greenland): ejnar mikkelsen this seven-man expedition was organised and led by ejnar mikkelsen, and had as its main aim the recovery of the lost diaries and journals of mylius erichsen and høeg-hagen, who had died with jørgen brønlund during danmark-ekspeditionen 1906–08. after a very difficult passage through the pack ice aboard the alabama, the expedition was forced to overwinter at kap sussi on the east coast of shannon (75°19´n). at the end of september 1909, a sledge journey was made northwards to lambert land (79°15´n), where jørgen brønlund’s body had been found in 1908, but no significant new documents were found on the body, and no traces of mylius-erichsen and høeg-hagen were found in the vicinity. in march 1910, a five-man sledge party embarked on a long journey northwards, crossing dove bugt and ascending onto the inland ice via the glacier storstrømmen. three men then explored northernmost dronning louise land (76°08´n) before returning to the alabama, while mikkelsen and iver p. iversen continued northwards across the margin of 28 the inland ice to the inner part of danmark fjord (80°34´n). from here they attempted to retrace mylius-erichsen’s route and located two cairn reports. returning home along the outer coast of kronprins christian land the two men met great difficulties, suffered from illness and hunger, and at one point abandoned their equipment and even their diaries to make a dash for danmark havn, where they arrived on 18 september. after a failed attempt to reach their abandoned equipment, they retreated southwards, only to find on reaching shannon on 25 november that the alabama had sunk. a house (subsequently known as alabama) had been built on shore, but there was no sign of their five companions, who had left for norway aboard the 7de juni in early august. in the spring of 1911, mikkelsen and iversen made a sledge trip northwards to recover their diaries, but it was not until the summer of 1912 that the two men were picked up from bass rock by the norwegian sealer sjøblomsten. the popular accounts of the expedition contain no new place names (e. mikkelsen 1913), but the official report including scientific observations (e. mik kelsen 1922) provides 23 new names, mostly given for members of the expedition committee, members of the expedition, and others who had assisted them. 1912–13 den danske ekspedition til dron ning louises land og tværsover nordgrøn lands indlandsis (the danish expedition to dronning louise land and across the inland ice): johan peter koch j.p. koch and alfred wegener, both of whom had been members of the 1906–08 danmark-ekspedi tionen, organised a four-man expedition whose principal aims were to study meteorological and glacial conditions at the margin of the inland ice (koch 1913; sigurðsson 1948; wegener 1961). a traverse of the main ice cap of iceland with their icelandic ponies was made to gain experience of travelling on ice, after which the expedition was transported to greenland aboard the godthaab, on loan from the danish government, arriving at danmark havn (76°46´n) on 23 july 1912. equipment un loaded at danmark havn and stormkap included a motorboat, 16 icelandic ponies, 20 tons of pony food and a house for overwintering. during the summer the expedition goods were transported overland and by motorboat, around and across dove bugt, as far as kap stop where further progress was halted until the fjord froze over in the autumn (fig. 10); several of their ponies were shot at 29 fig. 10. j.p. koch and alfred wegener with their two helpers (vigfus sigurðsson and lars larsen) disembarked from the godthaab on 23 july 1912 at danmark havn. their equipment included a motorboat, 16 icelandic ponies, 20 tons of pony food and a wintering house. from danmark havn the expedition travelled overland, and by motorboat, around and across dove bugt as far as kap stop, where they were forced to wait until the fjord ice froze. at kap stop messages were left in a bottle attached to a wooden pole anchored in a stone-filled barrel. the messages were recovered in 1989, although the barrel had been blown over by strong katabatic winds. kap stop. equipment was then sledged to the front of bredebræ, and about halfway across the glacier towards dronning louise land, at which point the winter house borg was erected. koch fell into a cre vasse on 5 november and broke a leg, but this healed well during the winter. in the spring of 1913, the journey was resumed with the remaining five ponies. dronning louise land was traversed from east to west via borgjøkelen, fari magsdalen and kursbræ, and several peaks in clud ing dronningestolen and kaldbakur were climbed. on 8 may the last nunatak was left behind and the crossing of the inland ice began, the west coast of greenland being reached north-east of prøven (72˚23´n) on 4 july. the majority of the 40 new place names found on the expedition maps are in dronning louise land; a large group of names commemorate members of danmark-ekpeditionen 1906–08, while others were given after danish localities, incidents on the journey, or the appearance of features. commercial activities, early mountaineering, geological mapping: 1919–1960 1919–24 a/s østgrønlandsk kompagni (east greenland company ltd.) østgrønlandsk kompagni was a danish trapping company founded in february 1919 on the initiative of former members of the 1906–08 danmark-ekspe ditionen. it was based on private capital, with some state assistance, but poor hunting and the loss of two ships in the ice led to its closure in 1924. the first group of 10 hunters sailed in 1919 aboard the dagny to the danmark havn region (76°46´n), and established hunting stations at danmark havn (danmarkshavnhuset) and hvalrosodden, with an oth er farther south at germania havn (74°32´n). the company eventually had 14 huts and stations between kap broer ruys in the south and hvalros odden in the north, including two taken over from the 1901 baldwin-ziegler expedition, and alabama on shannon built by the 1909–12 alabama expedition. in august 1920, the dagny was crushed in the ice off shannon, before it could reach the northern sta tions. the crew overwintered, but two died before the rescue ship teddy arrived in 1921. one of the hun ters, john tutein, was killed by a bear in february 1921. the teddy supplied the hunting stations in 1921, and also in 1922 and 1923. on the way home in 1923, a bad ice year, the teddy was crushed in the ice, but the 21 crew and hunters eventually reached land in the ammassalik region (bistrup 1924; dahl 1925; tutein 1945), and were picked up by the quest in 1924. in 1924 the godthaab was sent up to evacuate the remaining hunters from carlshavn, germania havn and sandodden, and the company suspended operations. descriptions of hunting with the company are given by lund (1926), and a general account of com pany activities by møller (1939) and lauritsen (1984). jennov (1945) records the total catch of the company’s hunters from 1919–24 as 679 foxes and 117 bears. numerous place names originated from the hun ters and the captains of the two ships. lists of huts and stations with their names are given by møller (1939) and p.s. mikkelsen (1994, 2008). most of these were named for their geographical locations, some for features and incidents, and a number for persons, including members of the board of directors of the company. møller’s account includes a sketch map from gustav thostrup’s 1921 logbook with about 20 names around eastern clavering ø. many of these names now have approved status. 1922–23 johan a. olsen expedition a seven-man norwegian expedition sailed to east greenland on the anni i, with the prime objectives of fox trapping and setting up a weather station at mygg bukta for the geofysisk institutt in tromsø. the station transmitted weather reports three times daily from 14 october 1922 until 15 august 1923, when the expedition began its homeward voyage. the anni i was lost with all hands, presumably crushed in the pack ice. 1923 was a bad ice year. 1924–25 foundation of scoresbysund harald olrik had proposed the foundation of a settlement in the unpopulated tracts of scoresby sund (70°–71°n) in 1911. the project was brought to fruition in 1924 due to the interest and influence of ejnar mikkelsen. the ‘scoresbysund-komiteen’ was founded on 24 march 1924 with ejnar mikkelsen as chairman, a post he was to hold for 40 years. an appeal to the danish public was immediately successful thanks to the support of valdemar galster, editor 30 of the ferslew press, and h.n. (hans niels) andersen of the østasiatisk kompagni that purchased a ship for the expedition, the fox ii that was renamed the grønland. the grønland left copenhagen on 10 july 1924 laden with building materials and provisions, made an easy passage of the ice belt and arrived off the mouth of scoresby sund on 24 july. at fox pynt near kap tobin the ship was caught in the ice and lost its rudder, an incident which led to immediate selection of a site nearby for the settlement without the plan ned preliminary reconnaissance (e. mikkelsen 1925). materials were unloaded at ferslew pynt, and the grønland returned home leaving behind a wintering party of seven, including three carpenters and three scientists. one of the latter, the geologist bjerring pedersen, died in july 1925, apparently of scurvy (bengtsson 1927). a large house was built at the present scoresby sund (the name of the settlement is spelt in danish in one word as ‘scoresbysund’, to distinguish it from the fjord known as scoresby sund) and small houses were built at kap stewart, kap hope and kap tobin for the greenlandic hunters and their families. about 16 names are associated with the colonisation expedition and reports of the overwintering scientists; some were given for expedition supporters and the ship, others record the bird and animal life. about 85 greenlanders arrived in 1925, the nucleus of what was to be a successful settlement (see also below). 1924–67 østgrønlandstraktaten (danish–nor we gian treaty on east greenland) the danish–norwegian treaty on east greenland (øst grønlandstraktaten) which came into effect in july 1924 gave both countries the right to engage in hunting, fishing and scientific activities in the uninhabited parts of east greenland, including the operation of meteorological stations. however, no agree ment was reached concerning sovereignty. the provisions of the treaty were exploited by both nations. denmark founded the new colony of scoresbysund, specifically allowed for by the treaty, and both norway and denmark developed trapping activities; norway re-opened the radio and weather station at myggbukta. danish scientific activities were initiated by lauge koch in 1926, the first of a succession of mainly geological expeditions under his leadership which continued until 1958. norway also embarked on scientific explorations, the norges svalbardog ishavsundersøkelser (nsiu – norwegian svalbardand arctic ocean survey) expeditions of 1929–33, but these were suspended when the dispute over the sovereignty of east greenland was determined in denmark’s favour by the court of international jus t ice at the haage in april 1933 (blom 1973; skar stein 2006). the treaty was to have lasted for 20 years, after which it could be terminated with two years notice. after the 1939–45 war, in which both danish and norwegian hunters had co-operated as members of nordøstgrønlands slædepatrulje, the treaty was ex tended (bruun 1966) and both danish and norwe gian fox-trapping activities were resumed. however, the value of fox skins had halved, and in practice the trapping companies were only able to exist with state subsidies. the danish state withdrew its subsidies to nanok in 1952, and norway similarly withdrew its subsidies to arktisk næringsdrift in 1959. after 1959 there was effectively no longer a norwegian presence in east greenland. denmark therefore took advantage of the termination clause of the treaty, and gave two years notice of its intentions in 1965. østgrøn landstraktaten was finally suspended on 9 july 1967; one of the principle arguments for the move was the need to establish a national park in north-east green land to protect its wildlife (bruun 1966). 1925–36 campagne du pourquoi pas? (greenland voyages of pourquoi pas?): jean-baptiste charcot the french polar explorer jean-baptiste charcot made numerous voyages to the arctic in his threemast barque pourquoi pas?, of which seven visited the scoresby sund region (charcot 1929, 1938; faure 1933). charcot was france’s leading polar explorer, the ‘father of french polar research’, and had earlier led two major expeditions to the antarctic in 1903– 05 and 1908–10 (malaurie 1989) (fig. 11). during his first visit to east greenland in 1925, to the newly founded settlement of scoresbysund (70°29´n), a short trip was made to nearby jameson land. in 1926 ejnar mikkelsen and ebbe munck travelled up as guests on the pourquoi pas? when charcot made a second visit to the scoresbysund settlement. the voyages between 1931 and 1933 were mainly concerned with the french polar station for the international polar year 1932–33 established at scores bysund. before leaving for home in 1932, the 31 pourquoi pas? visited the kap leslie area of milne land with lauge koch (see also fig. 71). charcot re turned in 1933 to pick up the international polar year wintering party, and the station buildings were handed over to the settlement. the pourquoi pas? also brought up the three-man ‘1933 cambridge east greenland expedition’ that worked in the hurry inlet area. charcot once again visited the kap leslie area. charcot returned to scoresbysund in 1934 and 1936, but on the voyage back to europe in 1936, the pourquoi pas? was wrecked on 15 september in a severe storm just after leaving reykjavik in iceland; only one crew member survived. about 20 names are linked with charcot’s expeditions, only one of which is commemorated on mod ern maps, a minor peak on milne land known as pourquoi pas tinde. the localities charcot gletscher and charcot havn, also located on milne land, were named subsequently by lauge koch’s expeditions. charcot’s place names are found in scientific reports of the work in the kap leslie region, and on a map of the area around scoresbysund (rothé 1941). 1925–present: scoresbysund / illoqqortoormiut [ittoqqortoormiit] the first party of greenlandic settlers, about 70 from ammassalik and 15 from west greenland, arrived at scoresbysund with the gustav holm (formerly the grønland, and originally the fox ii) on 1 sep tember 1925. different accounts give slightly different figures for the actual number of settlers. pho to graphs indicate there were a large proportion of children. the 15 from west greenland were henrik høegh (later colony manager) and the priest sejer abelsen, and their families. the first colony manager was johan petersen, former manager of the ammassalik colony for 30 years (nielsen 1957). the first few months were made difficult by an influenza epidemic, picked up when the ship called at iceland. everyone became ill, and three women, one man and a child died. by the end of the first year, however, 10 hunters had achieved a catch of 12 narwhales, 700–800 seals, 60 walruses, 115 bears and 75 foxes, and favourable hunting subsequently ensured the survival of the settlement (e. mikkelsen 1989). however, walrus were reported as rare after 1926. in 1926 the colony was reinforced by a family of 10 from west greenland, and in 1935 by a further 31 greenlanders from ammassalik (e. mikkelsen 1950). the greenlandic name for the settlement of sco res bysund started as igtorqortôrmît, which translates as ‘those that live at the place with one large house’. e. mikkelsen (1950) describes the large house as com prising living quarters for the families of the colony manager and the priest, which were separated by a small shop. when a church was built at the settlement in 1928, the priest had his own residence at tached to the church and there was also space for a school. the greenlanders lived at first in the villages of kap stewart, kap tobin and kap hope, near the best hunting grounds. a tendency for a concentration of the population at scoresbysund was later reported, 32 fig. 11. the french polar explorer jean-baptist charcot [1867–1936] made seven voyages to the scoresby sund region with his 3 mast ice strengthened barque pourquoi pas?. charcot was drowned on 16 september 1936, when pourquoi pas? was wrecked shortly after leaving iceland with the loss of 39 crew and scientists; only one man survived. photo: kindly supplied by emilie thomassot, © centre de recherches pétrographiques et géochimiques, nancy, france. allegedly due to the influence of the priest. kap stewart proved liable to heavy snow, and was abandoned in 1930. in 1947 two hunters with their families moved to a new settlement established west of kap brewster on the south side of scoresby sund. hunters also spent periods at sydkap in 1934–35, and a shop and store house were built there in 1946; however, this site has only occasionally been occupied. hunting huts have been built in several areas, including hurry inlet, steward ø, the coast of jameson land and the east coast of liverpool land. in 1928 scoresbysund was expanded with the addition of 10 houses, as well as the church noted above. a radio station was established by janus sør ensen in 1927. in 1932 the french expedition house, built for the international polar year, was taken over by the settlement, and used first as the telegraphist’s house, and later as a hospital. a new hospital was built in 1957 after a fire had destroyed the old buiding. during world war ii, american forces operated a weather station manned by 20–30 men in hvalros bugten nearby. a larger weather and radio station was established at kap tobin just south of scoresbysund in 1947, and closed down in 1980. the population of scoresbysund / illoqqor toor miut was 430 in 1983, with an additional 79 at the settlements at kap tobin and kap hope (statistisk årbog 1984), and in 2009 a total of 489 persons all in scoresbysund (statistisk årbog 2009). in 1983, there were 77 persons licensed as full-time hunters and 99 as part-time hunters. the yearly catch by registered hunters totaled about 6000 ringed seals, 50–70 polar bears, and smaller numbers of other seals, narwhales and walruses. the activities of greenpeace and bri gitte bardot have influenced the market for ringed seal skins since 1978, and as a result bear skins have provided an increased proportion of income. spring hunting for polar bears now ranges far afield, south along the blosseville kyst, north to daneborg, and westwards to gåsefjord. the greenlandic population has given numerous names to features in the vicinity of the settlements and the main hunting grounds. about 190 names were recorded by the 1955 geodætisk institut name registration, all of which were approved. the spelling of the greenlandic name for the settlement that began as igtorqortôrmît became ittoq qor toormiit in the east greenland dialect following the revision in spelling (see e.g. arke 2003). however, a west greenland dialect spelling illoqqortoormiut that had appeared in many ministry for greenland documents in the 1970s, was applied on official maps in 1995 for the town; however, the ‘ittoqqortoormiit’ spelling variation officially survives for several names derived from their proximity to the town (e.g. ittoq qortoormiit ilinnerat, ittoqqortoormiit kimmut kang er tivat, ittoqqortoormiit qinngerajivat). 33 28° 22°w 73°n69° fig. 12. while james wordie investigated a possible route to petermann bjerg (climbed during his 1929 expedition) during the 1926 cambridge east greenland expedition, other members of the party carried out geological and other investigations. this is part of the map surveyed by white (1927), showing the location of three of william scoresby jr.’s capes (scoresby 1823) that he was able to identify as mountains (mt. freycinet, mt. leitch, mt. laplace; now freycinet bjerg, leith bjerg, laplace bjerg). 1926 cambridge east greenland expedition: james mann wordie j.m. wordie led an eight-man expedition to east greenland in 1926, travelling aboard the heimland with lars jakobsen as captain. most of the scientists were from cambridge university in england. the expedition aims included surveying, archaeology and exploration of a route to the 2970 m high mountain of petermann bjerg (73°05´n) seen from a distance by karl koldewey’s 1869–70 expedition (wordie 1927). a similar expedition in 1923 on the smaller heimen had failed to reach the coast due to very bad ice conditions. the 1926 expedition left aberdeen on 30 june, stopped briefly at jan mayen, then made an easy passage of the ice belt to reach lille pendulum on 12 july. pendulum experiments were made on sabine ø (74°35´n), repeating sabine’s observations of 1823. during the summer, extensive surveying was carried out around the pendulum øer, the west side of clavering ø (where granta fjord was discovered), hold with hope and the interior of loch fyne (leading to the discovery of stordal), and along the outer poorly known coasts of geographical society ø and traill ø. from the inner part of kejser franz joseph fjord a route to petermann bjerg via ridderdal was explored, but the short time available prohibited an attempt on the peak. the heimland left the east greenland coast on 25 august after calling briefly at scoresbysund. in addition to the great improvements to existing charts in the coastal region, success was achieved in correctly placing many of the features named by william scoresby jr. in 1822 (fig. 12); many of his capes proved to be mountains standing well back from the coast (white 1927). about 30 new names were proposed for the coastal region and the area west of kjerulf fjord, some commemorating polar explorers, others cambridge locations and the general appearance of features. 1926–27 lauge koch’s geological expedition lauge koch’s east greenland expedition of 1926– 1927 comprised three geologists and two greenlandic dog-sledge drivers, and had as its object a general geological survey of the region north of scoresby sund (70°15´n). the danish geologist lauge koch [1892–1964] had already made his name as a member of knud rasmussen’s 2nd thule expedition, and especially for his geological and topographical map ping during his own ‘jubilæumsekspeditionen nord om grønland’ (jubilee expedition of north green land) 1920–23. the 1926–27 expedition was the first of a long series of east greenland geological expeditions led by lauge koch that were to continue until 1958. the expedition travelled to greenland with the gustav holm in july 1926. in august and september two geologists, alfred rosenkrantz and tom harris, worked in eastern jameson land (70°50´n), while koch organised construction of an expedition house in scoresbysund. in october koch made a sledge journey northwards to hold with hope via hurry inlet, kong oscar fjord and sofia sund, returning westward around ymer ø and retracing his outward track in november. between february and june 1927 koch made a long sledge journey to danmark havn (76°46´n). on the return journey the fjord system between 72° and 74°n was explored, and an unexpected extension of dusén fjord discovered. meanwhile rosenkrantz and harris had continued their work in jameson land, and also on eastern milne land. rosenkrantz made a journey to the interior of gåsefjord in 1927 to search for h.k.e. krueger, a german geologist erroneously supposed to have crossed the inland ice. the expedition returned to denmark aboard the gustav holm in august 1927. the main geological results of the expedition are described by koch (1929a, b, 1930a), and include a geo logical reconnaissance map of the region 70°–76°n. both alfred rosenkrantz, a geologist and palaeontologist based at the mineralogical museum in copenhagen, and tom harris, a palaeobotanist from cambridge university in england, carried out pioneer investigations of the mesozoic sedimentary rocks of the jameson land region. koch’s sledge journeys gave rise to about 12 place names, while the work of alfred rosenkrantz and tom harris gave rise to an additional 47 place names, mainly in jameson land and southern liverpool land. these were given for the shape and character of features, for geological associations such as finds of fossils, for animals, and for a few persons including their greenlandic assistants. many of these names first appeared on maps drawn by lauge koch (koch 1929a), and others in reports by harris (1931) and rosenkrantz (1932, 1934, 1942). 34 1926–28 foldvik expedition the norwegian foldvik expedition was the third to overwinter in east greenland, but broke new ground in adapting techniques of hunting used in spitsbergen and jan mayen to the larger greenland terrains. the practice of building numerous small huts over a wide area around a central station was subsequently fol lowed by all norwegian and danish hunting expeditions. the 1926–28 expedition comprised nils fold vik, hallvard devold and fritz øien, all tele graphists from the geofysisk institutt (geophysical institute) in tromsø, who with three hunters travelled to greenland in 1926 aboard the ringsel. two hunting stations were built, at revet (74°22´n) and near kap stosch (krogness; 74°03´n), and 17 huts in the sur rounding areas. hunting was carried out between kap bennet in the south and tyrolerfjord in the north, the catch including 287 foxes, 18 bears and seven wolves. the expedition returned to norway aboard the terningen in 1928. a short account of their work is given by foldvik (1933). 1927–28 scoresbysund seismic and radio station: janus sørensen following a short visit to scoresbysund (70°29´n) in 1926 to choose a site, janus sørensen returned in 1927 to erect a radio station and seismic station at the settlement. the latter operated until 1948, when it was moved to kap tobin. janus sørensen made sledge journeys around the coast of southern liverpool land, as a result of which a simple map was prepared that included several new names, including kap høegh, named after the colony manager (sørensen 1928). 1927–29 hird expedition: jonas karlsbak this six-man norwegian expedition led by jonas karlsbak took its name from the 49-foot fishing boat hird which carried it to greenland, and which sank in its winter harbour in the finsch øer (74°n) in august 1927. the expedition built three hunting sta tions, one at kap herschel, another on the south-east side of clavering ø (elvsborg), and the third on jackson ø; in addition seven huts were erected, of which five were on wollaston foreland. their catch amounted to 352 foxes and 42 bears. they returned home with the veslekari in 1929 (giæver 1939). 1927–29 alwin pedersen – scoresbysund as a follow up of his work in 1924–25 on the expedition that had founded scoresbysund, the german zoologist alwin pedersen organised an independent expedition to continue his studies. two years were spent at scoresbysund (70°29´n), during which he made a number of sledge journeys, one of them to the interior of nordvestfjord which led to the discovery of new arms of the fjord and the finding of polar bear dens (pedersen 1930). another trip took him south of scoresby sund as far as kap dalton. 1928–30 finn devold’s expedition a six-man norwegian hunting expedition led by finn devold sailed to east greenland in 1928 on the ter ningen, taking over the foldvik expedition terrain. a larger station was built at revet (74°22´n), and four new huts. their catch amounted to 346 foxes, 11 bears and 8 wolves (giæver 1939). the expedition returned to norway in 1930 with the vesle kari. 1929 cambridge east greenland expedition: james mann wordie wordie’s nine-man expedition from cambridge uni versity in england, had two prime aims: the ascent of petermann bjerg (2970 m; 73°05´n) and geological exploration. the heimland that had been used in 1926 was again chartered, captained by karl jakob sen, and departed from aberdeen on 2 july. however, ice conditions were severe, and the coast of east greenland was not reached until 4 august. from the inner end of kejser franz joseph fjord six of the party set off via ridderdal for what proved to be a successful first ascent of petermann bjerg, via ptar migan gletscher, across nordenskiöld gletscher and up disa gletscher. the summit of petermann bjerg was reached via the south-west ridge on 15 au gust (wordie 1930a, b). meanwhile two of the geologists carried out regional geological studies from the ship (wordie & whittard 1930; parkinson & whit tard 1931). the survey work of the expedition, much of it carried out by r.c. (cuthbert) wakefield and augu stine courtauld, was mainly around the head of kej ser franz joseph fjord and petermann bjerg, and most of the 20 new place names are in this region. a few elsewhere derive from the geological work. the expedition left the greenland coast on 25 august, again meeting difficult ice conditions which took them five days to clear. 35 1929–30 lauge koch’s geological expeditions lauge koch organised a summer expedition in 1929, financed largely by private contributions with the balance provided by the carlsberg foundation and rask-ørsted foundation; the ship godthaab was supplied by the danish state. the expedition numbered 22, including the ship’s crew, four geological parties and one botanical party. difficulties were experienced in penetrating the ice belt both on the way in and out. work was mainly carried out in the fjord region between 72°–75°n, with topographical surveying of parts of clavering ø, wollaston forland, hudson land and ymer ø (koch 1930b). for the 1930 summer expedition, koch secured passage on the godthaab, which was to visit east greenland on a danish navy inspection cruise. there were two geological, one zoological and one botanical parties on board. ice conditions created some difficulties, but work was carried out on clavering ø, and in parts of the kap stosch and moskusoksefjord areas. a brief description of the expedition is given in koch (1955 pp. 26–32). the summer expeditions of 1929 and 1930 visited the same general region and had many of the same participants. the majority of the c. 100 place names associated with these two expeditions are discussed by seidenfaden (1931), while others appear in the report of backlund (1932). about 45 names comme morate persons, including danish and swedish scientists, and members of j.m. wordie’s 1926 and 1929 expeditions. most others refer to incidents, or to characteristics of the features. 1929–33 norges svalbardog ishavs undersøkelser (norwegian svalbardand arctic ocean survey) norges svalbardog ishavsundersøkelser (nsiu) commenced scientific activities in east greenland in 1929 on the initiative of adolf hoel, a move coinciding with the foundation of arktisk næringsdrift a/s (see below) and the commencement of intensive land-based fox trapping. from 1929 to 1931 the scientific activities were on a modest scale, and included topographical surveying, oceanographical, botanical, zoological and geological investigations, mainly in the region between antarctic havn (72°n) in the south to wollaston forland (74°15´n) in the north. following the declaration of sovereignty over eirik raudes land (71°30´–75°40´n) by norway in 1931, the pace of activities was greatly increased. a 36 fig. 13. part of the 1932 map by norges svalbardog ishavsundersøkelser (nsiu) at a scale of 1:200 000. the map is part of the region that norwegian trappers called ‘eirik raudes land’, and the segment shown gives the location of myggbukta on the south coast of hold with hope (nsiu 1932a). major expedition sent up in 1932 with the polar bjørn included two aeroplanes to undertake aerial photography. the ruling of the court of international justice at the hague in april 1933 in denmark’s favour led to a reduction in activities. the nsiu scientific group in 1933 numbered nine and from 1934 scientific activities virtually ceased. however, nsiu continued to cooperate with arktisk næringsdrift in the dispatch of relief ships to serve the norwegian hunters, as well as supplying the telegraphists at myggbukta. the majority of place names associated with nsiu are found on map sheets published by nsiu at scales of 1:200 000 (fig. 13) and 1:1 million (nsiu 1932a, 1932b), the 1:100 000 topographic maps of lac mann (1937; fig. 14), and in expedition reports by orvin (1930, 1931) and nsiu (1937). lacmann lists the derivation of 299 new names appearing on the new maps, most of which were given for natural fea tures of the terrain (75), followed by norwegian place names (41), norwegian ships (32), hunters (30) and scientists involved in photogrammetric developments (26). only a selection of the many names used by nsiu has been officially approved for usage on danish maps of greenland, largely because of the nationalistic climate associated with the dispute over east greenland, and an impression that the namegiving was more prolific than necessary. however, a few of the nsiu names subsequently appeared on the united states air force 1:250 000 scale aeronautical charts published in the 1950s. 1929–41 østgrønlandsk fangstkompagni nanok a/s (east greenland trapping company nanok ltd.) østgrønlandsk fangstkompagni nanok (the east greenland trapping company nanok, commonly known as ‘nanok’) was founded in may 1929 on the basis of a plan by j.g. (johannes gerhardt) jennov, following several failed attempts to revive the old østgrønlandsk kompagni. the capital was secured by the support of several large danish companies. however, trapping was often poor, and nanok only survived with the assistance of the danish state, which provided free transport to and from green land, and the support of private funds, notably lau rits andersens fond, otto mønsteds fond, julius skrikes stiftelse, tuborg fondet and kaptain alf trol le og hustrus legat. the interest in the maintainance of danish hunting activities was largely a consequence of the challenge to danish sovereignty of east greenland by norway, and the necessity of competing with norwegian hunters. in 1929 nanok sent up 10 hunters with the bir gild, accompanied by jennov and the geologist richard bøgvad, but due to poor ice conditions only the southern hunting stations taken over from øst 37 fig. 14. segment of one of the 1:100 000 scale topographic maps published to support the norwegian claims to the part of east greenland they called ‘eirik raudes land’ (lacmann 1937). the excerpt shows the eastern part of vega sund that forms the south boundary of geographical society ø. most of the place names were new, and were not approved by the danish authorities. grønlandsk kompagni were occupied. transport to and from greenland was subsequently largely undertaken with the godthaab or the gustav holm, the two ships serving lauge koch’s geological expeditions. ice conditions often meant that stations in one or another area could not be reached, although j.g. jennov blamed the failure to relieve nanok’s stations in 1934 on lauge koch’s lack of interest in helping the danish hunters, a viewpoint unexpectedly supported by john giæver (lauritsen 1984). in 1935, the godt haab failed to reach the coast, but three hunters were evacuated by plane, and another four by the nor wegian sealer buskø. in 1937, the gustav holm became trapped by ice in scoresby sund, and no stations were reached. nanok had taken over 14 hunting stations from østgrønlandsk kompagni and built many new huts in the period 1930 to 1932. in 1932 the gefion was sent up to re-occupy the station at danmark havn, and a radio station was built at hvalrosodden (jennov 1935). following a fund-raising campaign numerous huts were built in 1938, and the company eventually had more than 60 huts between kap broer ruys (73°32´n) in the south and sælsøen (77°04´n) in the north. hunting success varied; 1931–32 and 1937–38 were reported as good trapping seasons, while trapping was poor in the 1934 to 1937 seasons. jennov (1945) reported the catch for the years 1929–38 as 1232 foxes and 67 bears. accounts of hunters’ experiences with nanok are given by drastrup (1932), hvidberg (1932), hansen (1939), kristoffersen (1969) and nyholm-poulsen (1985), and summaries of na nok’s activities by jennov (1935, 1939, 1945, 1953), lauritsen (1984) and p.s. mikkelsen (1994, 2008). operations were suspended in 1941 with the advent of war in europe, and the hunters returned home, moved to west greenland or north america, or joined nordøstgrønlands slædepatrulje. hunting was resumed in 1945. names originating from nanok are found in the descriptive, published accounts of the hunters, but notably in the maps and reports of jennov (1935, 1945) and the systematic descriptions of huts and stations by p.s. mikkelsen (1994, 2008). some names were officially approved, but others conflicted with the principles established by the then newly formed place name committee and were rejected. a large proportion of the names were suggested by j.g. jennov. 1929–42 arktisk næringsdrift a/s (arctic commercial enterprise ltd.) the norwegian trapping company arktisk nærings drift was founded in october 1929. following hall vard devold’s return from a private hunting ex pedition to east greenland, devold gained adolf hoel’s interest and support in greatly expanding norwegian hunting activities, while hoel saw the opportunity of developing nsiu scientific investigations (see above). arktisk næringsdrift began operations in 1929, and had hunters in east greenland continuously until 1942, and again from 1946 to 1959. the company had variable, often substantial, financial support from the norwegian state, and lesser amounts from the norwegian meteorological institute on whose behalf the myggbukta radio and weather station was operated from 1930. transport of hunters to and from greenland was undertaken by nsiu from 1929 to 1934, after which arktisk næ rings drift took over responsibility for ship charter for their own hunters (still in cooperation with nsiu), as well as those of private norwegian hunting expeditions. between 1929 and 1931, arktisk næringsdrift built 35 hunting huts between vega sund and mos kus oksefjord, and by 1938 with the other norwegian hunting expeditions had established 130 hunting huts and stations between canning land (71°41´n) in the south, and southern dove bugt in the north (76°15´n). on 29 june 1931, hallvard devold raised the norwegian flag at myggbukta and took possession of eirik raudes land, the region between 71°30´n and 75°40´n where norwegian hunters had been most active; this action was supported by norway who proclaimed annexation on 10 july 1931. the claim was contested by denmark, which appealed to the international court of justice at the hague; the case was decided in denmark’s favour on 5 april 1933, by a majority verdict (12 to 2). arktisk næringsdrift had 10 hunters in east green land from 1929 to 1931, and subsequently had 5–6 hunters active each year. many spent long periods in east greenland; gerhard antonsen wintered for a total of seven years at revet. norwegian hunters seem to have been generally more successful than their danish counterparts, arktisk næringsdrift reporting a catch of 3400 foxes and 26 bears between 1929 and 1938 (giæver 1939). in the season 1937–38 a single hunter at kap herschell caught a record 642 foxes. norwegian hunters are reported to have shot large 38 numbers of birds (schaanning 1933), including in the period 1928 to 1931 a total of 190 ravens, 40 snowy owls, 170 falcons (70 shot by finn devold at mygg bukta in 1928), 200 barnacle geese, 80 eider ducks, 65 red-throated divers and 2040 ptarmigans. supply ships visited the hunting stations every year; those used including the veslekari, polar bjørn, sælbarden, buskø and isbjørn. the supply ships occasionally carried small parties of tourists or sport hunters (munsterhjelm 1937). in spite of the outbreak of war in europe and norway’s capitulation, the veslekari was sent to east greenland in 1940 to relieve the norwegian hunting stations as usual. on its return voyage it was arrested by the fridthof nansen, a norwegian naval ship in the service of the allied forces, which also destroyed the radio facilities at myggbukta. in 1941 another supply vessel, the buskø, was arrested by the united states patrol boat northland. only three hunters wintered in 1941– 42, and in the summer of 1942 trapping operations were suspended. one hunter went to west greenland, another joined the us forces, while henry rudi remained in east greenland as a member of nordøst grønlands slædepatrulje. personal accounts of hunting activities and experiences in east greenland are given by giæver (1930, 1931), bang (1944), akre (1957) and winther (1970, 1980), and summaries of the work of arktisk næringsdrift and other hunting expeditions by giæver (1939) and lønø (1964). all hunting stations and huts had names, some incidental or commemorative, although many were known simply by their geographical location. a large number were known by different names at different times. the most exhaustive account of the stations and huts is that of p.s. mikkelsen (1994, 2008). 1930 robert a. bartlett east greenland expedition robert a. (bob) bartlett, the noted american skipper who captained the roosevelt during robert e. peary’s attempts on the north pole, made a journey to east greenland with his schooner effie m. mor ris sey in 1930, accompanied by the big-game hunter harry whitney. their main objective was to collect archaeological and anthropological specimens for the museum of the american indian, heye foundation (now part of the smithsonian institution). the expedition visited the coastal region between 74° and 76°50´n, kap bismarck being the northernmost point reached. archaeological excavations were made at kap david gray and eskimonæs (bartlett & bird 1931; bartlett 1934). 1930–31 constantin dumbrava’s scoresby sund expedition having spent several years in the ammassalik / tasii laq region, the rumanian scientist constantin dum brava moved his area of interest to the scoresby sund region, in defiance of the wishes of the danish authorities. the norwegian sealer grande, cap tained by bernt heide, had disembarked dumbrava with his equipment on the east side of hurry inlet in the summer of 1930; dumbrava built a house and made meteorological observations. the next year the godthaab was diverted to pick him up and extradite him to europe. his visit gave rise to use of three place names: dumbrava, dumbravap imia and dumbrava kangileqitaa, all of which were incorrectly spelt dombrava for many years. [place name committee archive.] 1930–31 deutsche grönland-expedition (german greenland expedition): alfred l. wegener the main 19-man party of alfred wegener’s expedition to undertake a systematic study of the greenland inland ice and its climate sailed to west greenland. the expedition ascended the ice cap using primitive tracked vehicles, and established the eismitte station. wegener died while attempting to return from eismitte to the coast of west greenland in november 1930 (wegener 1932, 1935). a three-man party led by walther kapp travelled to the scoresby sund region of east greenland in july 1930 aboard the gertrud rask, to establish wege ner’s eastern land station that was to carry out complimentary meteorological observations. initially studies were undertaken around the town of scores bysund, but in early september the party moved with the help of greenlanders to the west coast of jameson land where wegener’s oststation was established south of the present gurreholm (wege ner 1932, 1935). the party sledged back to scores bysund in may 1931, and in july sailed back to europe aboard the gertrud rask. only three names in the scoresby sund region are associated with this expedition, including one given by the greenlanders, tyskit nunaat. 39 1930–32 møre grønlands ekspedition (møre greenland expedition) this norwegian six-man hunting expedition was led by jonas karlsbak, and included four members who had previously hunted with the hird expedition. they travelled up in 1930 with the veslekari. three of the hunters opened up new terrain on the south side of kong oscar fjord with main stations at ant arctichavn and kap peterséns, and built twelve new huts between canning land and alpefjord. in au tumn 1931, one of the hunters, knut røbek, fell through the fjord ice and drowned. two men return ed home in 1931 because of illness, and the others in 1932 aboard the polarbjørn (giæver 1939; p.s. mikkelsen 1994). 1931 louise a. boyd’s arctic expedition this was louise boyd’s third arctic expedition, but the first to visit east greenland; the earlier expeditions were to franz josef land in 1926, and to spits bergen and franz josef land in 1928. louise arner boyd [1887–1972] had inherited her father’s considerable fortune in 1920, and her independent and adventurous spirit led to her becoming involved in arctic exploration. her 1931 east greenland expedition was primarily a photographic reconnaissance in preparation for the more ambitious 1933 expedition. the norwegian sealer veslekari was chartered, and in the course of the summer visited every fjord and sound between 72° and 74°n. the inner part of is fjord was visited for the first time and gerard de geer gletscher discovered, and from the south end of kjerulf fjord a new route to hisinger gletscher was explored and mapped. alpefjord and röhss fjord were also penetrated to their inner ends (anrick 1932; boyd 1932). the passengers included the big-game hunter harry whitney. a small group of names are associated with the expedition. 1931 von gronau’s flight over the inland ice the german aviator, wolfgang von gronau, with three companions made a pioneer flight in august 1931 from europe to north america in a dornier seaplane, ‘grönland-wal’, which included a crossing of the greenland inland ice from scoresbysund to maniit soq/sukkertoppen (gronau 1933). after taking off from scoresbysund strong winds were encountered in the inner part of the fjords. a diversion was made southwards to gain altitude, in the process flying over unexplored mountains south of scoresby sund; one group of these mountains now bears the name gro nau nunatakker. 1931 høygaard & mehren expedition the norwegians arne høygaard and martin mehren made a crossing of the inland ice from west to east in july and august 1931. on 6 august they sighted the first nunataks of east greenland at about 73°30´n, and during the next ten days made their way through the unexplored glaciers and nunataks between 73°30´ and 74°10´n, eventually reaching northern strind berg land, and via waltershausen gletscher the west coast of nordfjord. the return to norway was made with the polarbjørn (høygaard & mehren 1931). of the 14 new names recorded, nine commemorate norwegians who had assisted them, or had con nec tions with arctic whaling or exploration. other names were given for the appearance of features. 1931 norcross-bartlett expedition to the greenland sea robert a. bartlett again visited east greenland with his schooner effie m. morrissey, this time in company with arthur d. norcross. the aims were similar to his 1930 voyage, to make collections for the smith sonian institute, the american museum of natural history and the heye foundation. ice conditions off east greenland were very difficult, and the ship was trapped for 37 days before land was reached at clave ring ø. visits were made to kap stosch, shannon and a few other localities (bartlett 1934). 1931–34 treårsekspeditionen til christian x’s land (the three-year expedition to east greenland): lauge koch treårsekspeditionen was the largest and most com prehensive expedition hitherto sent to east green land by denmark. the financial support came largely from the carlsberg foundation and from private contri butions, while government support was in the form of transport in the ships gustav holm and godthaab and seaplanes borrowed from the danish navy. topographical surveying was entrusted to the geodetic institute (geodætisk institut). the expedition was to extend over four summers and three winters, the scientists wintering in specially built stations. the specific tasks of the expedition includ40 ed preparation of topographic maps of the region 72°–76°n, together with geological, zoological, bota nical, archaeological and hydrographical studies in the same region. general accounts of the expedition are given by thorson (1937) and koch (1955). lauge koch was empowered as the danish police authority in east greenland pending the verdict on sovereignty of east greenland by the international court of justice at the hague. after the decision in favour of denmark, ejnar mikkelsen was appointed inspector for east greenland under the authority of grønlands styrelse (the greenland administration) although in practice lauge koch continued to represent police authority in east greenland during his expeditions until 1939. the 1931 expedition numbered 65, including 22 scientists and their assistants. the principal task of the first year was construction of the two main win tering stations at eskimonæs and on ella ø, and two smaller houses at nordfjord and kap brown (see also fig. 40). scientific work of all kinds was commenced, but was not extensive during the summer because of difficult ice conditions and house-building. geolo g i cal work was carried out mainly on clavering ø, ymer ø, traill ø and hochstetter forland. ten scientists overwintered in 1931–32, and a great deal of scientific work was carried out during autumn and spring sledge journeys. the 1932 expedition numbered 95, including 37 scientists and their assistants. two sea-planes were borrowed from the danish navy, one carried up aboard the gustav holm, and the second brought up on the french ship pourquoi pas?. the danish army flying corps provided four aerial photographers. the air support meant a considerable increase in the effectiveness of the cartographic work, with aerial photography supporting the ground trigonometrical surveys. on the basis of reconnaissance flights a working chart was prepared of the region from 70° to 77°n and was published in 1932 at a scale of 1:1 million (geodætisk institut 1932); it included many hitherto unexplored areas along the margin of the inland ice, (fig. 15). a new house (kulhus) was built during the summer on hochstetter forland. scien 41 fig. 15. the 1932 1:1 million scale geodætisk institut published map was drawn during the 1931–34 treårsekspedition by lauge koch (geodætisk institut 1932), and is partly based on aerial observations. following public criticism of lauche koch’s naming policy the glacier names gerda gl. and (a)nna sten gl. indicated at the top were not approved. 24˚w 22˚w 20˚w 74˚n tific studies were carried out between hochstetter forland in the north and traill ø in the south. zoological and hydrographical investigations based on the godthaab were carried out in most of the fjord system from 72° to 74°n. archaeological studies were made on the thule cul ture sites on clavering ø (dødemandsbugten), and in the district around ella ø. icelandic ponies were used with some success for the transport of camp equipment and geological samples. weather and ice conditions were more favorable than in 1931. twelve scien tists overwintered in 1932–33. the summer of 1933 saw the culmination of the expedition, which numbered 109, of whom half were scientists. weather and ice conditions were very favourable, and in august the gustav holm reached as far north as the norske øer off lambert land (77°n), from where reconnaissance flights were made northwards to peary land. aerial photography was undertaken throughout the region between 72° and 76°n, and the groundtrigonometrical survey was completed. geological studies extended from liver pool land in the south to skærfjorden in the north, and westwards to the innermost parts of the fjord systems. a mining camp was established on clavering ø to investigate a mineralised dyke with a conspicuous gossan; the ‘gold mine’ was found to comprise 90% pyrite and trace amounts of gold and silver. the godthaab undertook zoological and hydrographical studies in the scoresby sund fjord system. eleven ponies were used for transport, mainly to supply the mining camp. seven scientists overwintered in 1933–34. the l934 expedition numbered only 65, including 31 scientists and assistants, and had only one ship, gustav holm, and one sea-plane. the main work of the summer was geological, including work in the coastal region between canning land and hudson land, while inland eugène wegmann’s party reached cecilia nunatak and helge g. backlund’s party investigated the inner scoresby sund fjord system. poor weather and bad ice conditions hindered activities, and in particular prevented planned relief and transport of supplies to hunters of the nanok company. the intensive scientific activity over most of the region north of latitude 70°n, reaching many previously unexplored regions, gave rise to the introduction of a very large number of new place names. more than 480 names are credited to members of the 1931–34 expedition. most of the names can be attributed to particular expedition members as summa r ised below. [place name committee archive.] hermann aldinger visited the kap leslie region of eastern milne land in 1933 and introduced 22 new names. they were given chiefly for geological characteristics and geographical location, except for four names for french scientists who had worked in the same region with j.-b. charcot’s expeditions (al din ger 1935). helge g. backlund joined the expedition in the years 1932–34, although the majority of his names derive from his 1933 explorations of liverpool land. his 16 names were given for geological phenomena and other natural features, with some personal names. many other names proposed by backlund have not appeared in print, the alternative suggestions by laurits bruhn (mainly given after danish localities) being preferred by the place name com mittee, although as backlund pointed out they have no resemblance to their namesakes in denmark. laurits bruhn was a member of the geodetic in stitute surveying party in 1932 and 1933, when he was mainly at work in the scoresby sund region. of his 98 recorded names, many were named after danish localities, in particular the fjords of eastern liverpool land and the rivers of jameson land (map 4). other names were given for the appearance of features, a few for characters in fairy tales, while the precipitous cliffs of the volquaart boon kyst and its hinterland were given names with an element of fantasy. the ella ø overwintering party of 1931–32, that included ole simonsen, arne noe-nygaard and gun nar thorson, is credited with 10 names in the vicinity of the station, mainly named after the appearance of features and different animals. hans frebold undertook geological work in wol la ston forland and hochstetter forland in 1931. his 13 names were mainly given for geological features (frebold 1932, 1935). lauge koch was largely responsible for the numerous new names that appeared on the 1932 edition of the geodetic institute (geodætisk institut) 1:1 mil lion scale topographical map of the region from 70° to 77°n (geodætisk institut 1932), which included ex tensive, previously unmapped regions. many of his 59 names were given for danish politicians, army and navy officers and scientists who had assisted his expeditions. some were given for british and american scientists. the overwintering parties at kulhus in 1932–33 and eskimonæs from 1931 to 1934 were credited with 45 names in the region from 74° to 76°n. they have a 42 variety of origins, including their geographical location, size, shape and colour, while a few derive from incidents and from norse mythology. david malmquist in company with thorvald sør ensen reached 77°n aboard the gustav holm in 1933. their exploration and mapping of the skær fjorden area gave rise to 23 names, given for the size and shape of features, incidents, and for family mem bers and friends (seidenfaden & sørensen 1937). eigil nielsen, a vertebrate palaeontologist, gave 25 names to features in the vicinity of kap stosch, arising from his work in 1932 and 1933. they record a mixture of geological features, shape and colour, together with four greenlandic names (nielsen 1935). arne noe-nygaard and the swedish palaeontologist gunnar säve-söderbergh gave seven names to fea tures in north-east clavering ø deriving from their joint work in 1931 (noe-nygaard & sävesöder bergh 1932). they were given for geological characteristics, shape and colour. a further 35 names originate from noe-nyegaard’s work in canning land in the years 1931 to 1934, often in association with other geologists (noe-nygaard 1934). these names were given mostly for natural features of the localities, or for existing named features nearby. some features were named after notable scientists. gunnar säve-söderbergh studied late palaeozoic stratigraphy and palaeontology between jameson land and clavering ø from 1931 to 1934. most of his 34 names were given for features in gauss halvø, and derive from geological characteristics, notable geologists, and girls’ names (säve-söderbergh 1932, 1933, 1934, 1937). ole simonsen was a member of the geodætisk institut (geodetic institute) surveying party from 1931 to 1933, and is credited with 52 names from parts of andrée land, frænkel land, suess land, nat horst land, traill ø and the stauning alper. many of them were named after danish place names, while others result from incidents, or derive from the ap pearance of features. three were given for his green landic assistants. ragnar spärck and gunnar thorson were en gaged in marine zoological studies aboard godthaab in 1932 and 1933 in the fjord systems between 72° and 74°n. they proposed l4 names, mostly given for natural features of the localities. the swiss geologist eugène wegmann carried out work in the inner parts of the fjord system from 1932 to 1934, including the first exploration of the interior parts of suess land, gletscherland and lyell land, and the first visit to cecilia nunatak. about 60 names have been recorded, the great majority given for swiss localities, and a number for french and swiss scientists (wegmann 1935). a further 30 names arose during the expedition, but cannot be credited with any certainty to particular members. some are botanical localities apparently first used by gelting (1934), and others derive from a journey along the margin of the inland ice by th. sørensen and others in 1932 (koch 1940). [place name committee archive.] 1932 østgrønlandsk fangstkompagni nanok (east greenland trapping company nanok): gefion expedition j.g. (johannes gerhardt) jennov led an expedition in the gefion in 1932 with the objective of re-occupying the danmark havn trapping station and establishing and extending danish hunting activities in the dove bugt region (75°–77°n; jennov 1935). a radio station was established at hvalrosodden. a number of new names appear on the map published in jen nov’s account of the voyage, but very few of them were officially approved in spite of repeated applications to the place name committee. [place name committee archive.] 1932 scoresbysund committee second east greenland expedition: ejnar mikkelsen ejnar mikkelsen, chairman of the scoresbysund com mittee for more than 40 years, was leader of this expedition to the relatively poorly known coastal region south of scoresby sund. the aims were in part scientific, and in part to erect houses at suitable locations to enable communication between the settlements at ammassalik / tasiilak and those of scores bysund / illoqqortoormiut (ittoqqortoormiit). the expedition included british and danish scientists and sailed from copenhagen on 22 june aboard the søkongen, reaching the greenland coast at kap dalton on 10 july (e. mikkelsen 1933). scientific work was begun here and extended progressively southwards, detailed work being carried out in the kangerlussuaq region (68°–68°30´n). the expedition left ammassalik for copenhagen on 10 september. only one new place name is recorded north of latitude 69°n, høst havn, a bay near kap barclay. 43 1932 skaun & welde – ‘dagsposten’ expedition sigurd skaun and harald welde visited east green land with the support of the norwegian newspaper ‘dagsposten’ and adolf hoel, to investigate supposed columns of smoke seen by arne høygaard and martin mehren in 1931 on the east side of waltershausen gletscher. they travelled to greenland with the polar bjørn, and were landed at kap bull at the mouth of moskusoksefjord. a three week journey in difficult terrain in western hudson land and ole rømer land yielded no evidence of volcanic activity or hot springs (skaun 1932). their explorations gave rise to 12 new names, eight of which have come into general use with approved status. they returned home with the polarbjørn. in 1952 further sightings of ‘smoke’ in this region were reported by charles swithinbank and others aboard the polarbjørn, who were convinced that it was due to volcanic activity; this gave rise to reports in the ‘new yorker’ and norwegian newspapers. an unpublished letter by lauge koch, dated 1953, states that he is familiar with the ‘smoke’ in this region that consists of clouds of dust derived from driedout silt deposits on the floor of an ice-dammed lake beside waltershausen glet scher, periodically disturbed by strong winds [geus archive.] 1932–33 7th thule expedition: knud rasmussen knud rasmussen [1879–1933] was a danish –green landic polar explorer and anthropologist, most noted for his ‘thule expeditions’, that take their name from the trading station he established with peter freu chen in north-west greenland in 1910. the 7th thule expedition, the last of knud rasmussen’s thule expeditions, involved major scientific investigations along the south-east coast of greenland from kap farvel in the south to kangerlussuaq (68°30´n) in the north. emphasis was placed on surveying, and a sea-plane was supplied by the danish navy to undertake aerial photography. geological, archaeological, botanical and zoological studies were also prominent, and in 1933 knud rasmussen was notably involved in the production of a cinematographic record of greenlandic inuit life. almost all the work of the expedition was south of 69°n, but some of the aerial photography extended into the almost unknown region of high mountains and glaciers between kangerlussuaq (68°30´n) and scoresby sund, a region that figures prominently in official reports of the expedition as knud rasmussens land (gabel-jørgensen 1940). rasmussen had sailed along the blosseville kyst in august 1933 aboard the kivioq on the way to visit scoresbysund, returning to ammassalik by the same route. knud rasmussens land was the official name of the region between kangerlussuaq and scoresby sund (68°30´–70°n) from 1936 to 1953, but was then abandoned when the name was transferred at the suggestion of eske bruun (head of grønlands styrelse – the greenland ad mi n i stration) to cover much of western north green land, explored by knud rasmussen during the 1st and 2nd thule expeditions. the official ‘knud rasmussen land’ is very rarely used as a place name due to the very broad region which it now covers. however, the region between kangerlussuaq (68°30´n) and scores by sund is still commonly referred to as knud ras mussens land, especially in mountaineering lite r a ture. 1932–33 international polar year: j.-b. charcot jean-baptiste charcot had selected the site for a french scientific station at scoresbysund in 1931. in 1932, the pourquoi pas? and the french icebreaker pollux carried materials and personnel to set up the station, which comprised a main building ker dou mer and a smaller house ker virginia. the station was manned until the summer of 1933 (rothé 1941). elsewhere in east greenland the norwegian weath er stations at myggbukta and jónsbu took part in the international polar year project. 1932–34 sigurd tolløfsen’s expedition a norwegian six-man hunting expedition led by sigurd tolløfsen travelled to east greenland together with john giæver’s expedition aboard the isbjørn in 1932. tolløfsen’s party used the arktisk næringsdrift terrain between revet and godthåb gulf (74°– 74°30´n), and the so-called sunnmøre terrain from jackson ø to kuhn ø (73°50´–75°n). the expedition expanded the northern terrain with a new station, sigurdsheim, and six new huts. one of the hunters, arnljot tolløfsen, was drowned between loch fyne and kap herschel, and the remaining five went home with the nsiu relief ship sælbarden in 1934 (giæver 1939). 44 1932–34 helge ingstad’s expedition this six-man expedition was led by helge ingstad, a norwegian writer and lawyer who had been appointed sysselmann (= governor) of eirik raudes land following norway’s declaration of sovereignty over parts of east greenland in 1931. the expedition went up with the polarbjørn and took over the territory on the south side of kong oscar fjord. several huts were built, and a number of sledge journeys made, including one in the spring of 1933 across jameson land to the interior of nordvestfjord (ingstad 1935, 1937). after news that norway had lost the court case in the hague was received, ingstad returned home in 1933 with the polarbjørn, while the remainder of the expedition returned to norway with the sæl barden in 1934. 1932–34 john giæver’s expedition john giæver’s six-man hunting expedition travelled up with tolløfsen’s expedition on the isbjørn. they established the trapping and radio station jónsbu, which operated from 1932 to 1934, and two other hunting stations north of ardencaple fjord (otto strand and olestua). eighteen hunting huts were built between the south coast of ardencaple fjord and kap niels (75°–76°24´n), including two inland by large lakes, together representing a considerable expansion in the range of norwegian hunting activities. the expedition returned home with the sælbarden in 1934 (giæver 1939). 1933 louise boyd’s arctic expedition louise a. boyd’s fourth arctic expedition was organised with the cooperation and assistance of the american geographical society, and included five scientists: two surveyors, a physiographer, a geologist and a botanist. the botanist developed appendicitis and returned home without reaching greenland. the veslekari, captained by johan olsen, was the expedition ship and left norway on 28 june for jan mayen and greenland. hold with hope was reached on 13 july after an easy passage through the ice. nearly all the fjords from 72°30´ to 74°n were visited, and the expedition departed from mackenzie bugt on 9 sep tember (boyd 1935). louise boyd continued, during this voyage, her primary interest of making a photographic record of arctic scenery. for the 1933 voyage the veslekari had been fitted with an echo sounder, and profiles were successfully made in all the fjords, as well as on the atlantic crossing. knækdalen (gregory valley) was discovered and explored for the first time, and a photogrammetric map was made of the valley, as well as detailed maps of glaciers in knækdalen and on louise boyd land. in the course of geological studies noel e. odell ascended a number of mountains around knækdalen and in other areas (odell 1934a, b, 1937a, b, 1939, 1943, 1944). tidal guages set up at two localities gave useful information. about 20 new names are associated with the expedition, nearly all arising from the exploration of knækdalen (boyd 1935), and were given mainly for the appearance of features. 1933 charles lindbergh’s flight across greenland the american aviator charles lindbergh and his wife crossed the greenland inland ice from west to east on 4 august in their lockheed sirius monoplane ‘ting mis sartoq’ as part of a six month series of flights which took them around much of the north atlantic ocean. lauge koch provided them with weather reports, and they landed at ella ø, subsequently visiting eskimonæs on 5 august. on 6 august the lind berghs flew south to ammassalik, with instructions from koch to pay particular attention to the high mountains south of scoresby sund. they re-crossed the inland ice westwards to nuuk (then known as godthåb), then rounded the south coast of green land and flew back to ammassalik. at ammassalik they were entertained by knud rasmussen on 13 august, before departing the next day for iceland (lindbergh 1934). lauge koch subsequently named a group of nunataks south of scoresby sund after lind bergh. 1933 cambridge expedition to east greenland g.c.l. (colin) bertram, david lack and brian b. roberts, scientists based at cambridge university (england), travelled to east greenland in 1933 as guests of j.-b. charcot aboard the pourquoi pas?. zoological and ornithological studies were made around the inner part of hurry inlet (roberts 1935). most of their place names were adopted from the work of alfred rosenkrantz and tom harris, al though several were misplaced on their maps. 45 1933 john k. howard expedition to east greenland the american john k. howard visited east green land in august with the nordkap ii. a small geological party disembarked on western ymer ø (73°20´n), and their work gave rise to five new names (cleaves & fox 1935). two of their names were brought into general use by the next geologist to undertake systematic work in the area (eha 1953). 1934 count leonardo bonzi spedizione italiana (italian climbing expedition) a five-man italian climbing expedition led by leo nardo bonzi had intended to make an attempt on the watkins bjerge (69°n) from the blosseville kyst. however, the expedition ran into difficult ice conditions in their small icelandic boat njall, and turned their attention instead to the unexplored mountains behind volquart boon kyst (70°n) on the south side of scoresby sund. between 22 and 29 august parties explored and climbed a number of mountains and glaciers over an e–w distance of 35 km. thirteen names, nearly all with italian connections, were bestowed on a variety of features. bonzi’s (1935, 1936) sketch map proved difficult to reconcile with existing maps, and only three of his names were later adopted officially – savoia halvø, milano gletscher and roma gletscher. however, all bonzi’s peaks have since been identified on modern maps (fantin 1969). ice conditions de layed departure, and the expedition did not leave the greenland coast until 7 september. 1934 alfred rosenkrantz expedition to scoresby sund the danish geologist alfred rosenkrantz spent the summer in the scoresby sund region studying jurassic stratigraphy, assisted by greenlanders from scores bysund (rosenkrantz 1942). the greenlanders sub sequently gave the name ilimananngip nunaa to two of the areas where rosenkrantz worked, around kap leslie and around rødeelv in eastern jameson land. ‘ilimananngip’ translates roughly as ‘one does not expect anything from him’, implying that rosen krantz was not a generous employer; however alfred rosenkrantz was noted for his good relationships with the greenlandic members of his west greenland expeditions (niels henriksen, personal communication 2010). 1934 british trans-greenland expedition: martin lindsay martin lindsay led a three-man expedition to investigate the mountainous region south of scoresby sund in 1934, approaching the area after crossing the inland ice from west greenland by dog sledge. from the area of the gronau nunatakker the expedition traversed south-west around the head of kanger lussuaq (68°30´n), and eventually reached ammas salik. the expedition sailed back to europe with the jacinth (lindsay 1935). only a short time was spent north of latitude 69°n, and only three names are relevant to this ac count; two of these, prinsen af wales bjerge and grønlands styrelse gletscher, are approved. 1934–37 suløya grønlands ekspedition (suløya greenland expedition) this four-man norwegian hunting expedition in cluded two of the pioneers from the hird expedition, hermann andresen and peder sulebak. the group travelled up with the sælbarden, and hunted in two parties of two, on the south side of kong oscar fjord (72°n) and on wollaston forland (74°20´n). two men travelled home in 1936, and the other two in 1937 (giæver 1939). 1935 anglo-danish expedition to east greenland augustine courtauld and lawrence r. wager joined forces in 1935 for a summer expedition based at kangerlussuaq (68°30´n), with the primary aim of an ascent of the highest summit of the watkins bjerge. the 14-strong party included a danish archaeological group (eigil knuth, helge larsen and ebbe munck) as well as four wives of expedition members. on the way to kangerlussuaq the quest picked up two greenlandic families who were to experiment with hunting. in august 1935, a six-man climbing party, which included courtauld, wager and munck, embarked on the successful ascent of gunnbjørn fjeld (3694 m), the highest peak of the watkins bjerge, and the highest summit in greenland; a 190 km round trip via sorgenfri gletscher and christian iv gletscher (cour tauld 1936; longland 1936; munck 1957a, 1957b). the main peak lies south of latitude 69°n, but two new names given during this venture lie north of 69°n, guiden and ismågen. 46 the quest left kangerlussuaq on 29 august, leaving behind seven members who were to continue work as the 1935–36 british east greenland expedition. 1935–36 british east greenland expedition: lawrence r. wager this was a continuation of the 1935 anglo-danish expedition to east greenland and was made up of a party of seven led by lawrence r. wager, supported by a group of 14 greenlanders. the greater part of the work of the expedition was geological, and was car ried out south of latitude 69°n. wager discovered the ‘skaergaard’ intrusion, possibly the best known layered igneous intrusion in the world (wager 1937, 1947), that has subsequently been intensively studied by geologists and prospecting companies. two sledge journeys penetrated north of 69°n, one in the spring of 1936 up frederiksborg gletscher to gronau nuna takker and seward plateau, and the second in the summer of 1936 up frederiksborg gletscher, west of prinsen af wales bjerge, and south around the head of kangerlussuaq. the party returned to europe in late august aboard the seleis. these explorations gave rise to eight place names north of 69°n, and many more to the south outside the scope of this account. 1936 alfred rosenkrantz expedition to scoresby sund alfred rosenkrantz again spent a summer in east greenland studying jurassic stratigraphy, mainly in the area north of kap hope (rosenkrantz 1942). he was assisted by greenlanders from scoresbysund, and the expedition was made possible by financial sup port from the carlsberg foundation. 1936–37 quest expedition: gaston micard count gaston micard hired the quest, captained by ludolf schelderup, for a trip to east greenland, with the quest overwintering at the mouth of loch fyne (74°n). micard made use of norwegian hunting huts in loch fyne, and also built three new huts, later taken over by arktisk næringsdrift. two of the crew, willie knutsen and karl nicolaisen wintered at kap stosch (knutsen 1949). the crew of the quest caught 162 foxes. at the end of july 1937, the quest returned to europe, making short stops at scores by sund and ammassalik on the way. 1936–38 bird & bird ornithological expedition edward and charles bird spent respectively one and two years at myggbukta and peters bugt making ornithological studies (bird & bird 1941). transport and other facilities were provided by nsiu (norges svalbardog ishavsundersøkelser) and arktisk næ rings drift. 1936–38 two-year expedition: lauge koch this expedition, which had almost entirely geological objectives, was to last for three summers and two winters. each summer expedition was ship-based, with up to seven motor boats providing local transport, and in 1938 a sea-plane was used for aerial re connaissance. ponies were used extensively for trans port in jameson land. large wintering parties extended the field season using dog sledges for spring geological exploration. the expedition was financed in part by private contributions, the balance and loan of the ship being provided by the danish state (koch 1955). 1936 – the gustav holm carried 47 men to east greenland, reaching scoresbysund on 23 july. it was an exceptionally favourable ice year, no pack ice being encountered either on the voyage out or the voyage home. five geological teams were at work mainly between latitudes 71° and 74°n, including parts of gauss halvø, kap stosch, ella ø, traill ø and nat horst fjord. fourteen men wintered at the stations ella ø and eskimonæs. 1937 – ice conditions proved extremely difficult this year. one of the main objectives was the erection of a new wintering station, planned to be placed in nathorst fjord, but the gustav holm could not reach the area because of pack ice, and the new station gurreholm was built instead in western jame son land, near the mouth of schuchert dal. ice pre vented the relief of the northern wintering stations, with the result that the scientists who had intended to return home were forced to overwinter for a further year. eight geological, one zoological and one botanical team were at work during the summer in parts of hold with hope, the giesecke bjerge and jameson land (71°–74°n). twenty-three men overwintered at four stations. 1938 – the godthaab was expedition ship, and carried one additional geological party to greenland to join those already in the field. ice conditions again proved difficult, although not as bad as 1937. work 47 was carried out in hudson land, the giesecke bjerge, jameson land and scoresby land (71°–74°n). only two members overwintered, both returning home in 1939. hans stauber spent the entire period 1936–38 in greenland, wintering at ella ø and gurreholm, and working on traill ø, in scoresby land and jameson land. of the 21 place names that he proposed several had geological connections, some were derived from existing names, a few record incidents during the ex pedition and others commemorate swiss geologists. wolf maync and andreas vischer also spent 1936– 38 in greenland, wintering at eskimonæs. they gave 33 names to features on gauss halvø, the giesecke bjerge, wollaston forland and kuhn ø; these record geological associations, the appearance of features, or commemorate swiss localities and scientists. how ever, a further group of suggested names were considered unsuitable by the place name committee that proposed alternative names. un for tunately more than 30 unapproved names were used in their publications, and some of these have subsequently come into use as type localities of geological formations. heinrich bütler worked in the summers of 1936 and 1938 in hudson land and ole rømer land. most of his proposed names were given for swiss localities, swiss geologists, or for characteristics of the features. [place name committee archive.] 1937 louise a. boyd’s arctic expedition louise boyd once again chartered the veslekari, cap tained by johan olsen, for a voyage to east greenland and spitsbergen. scientific staff included two geologists, a botanist, a surveyor and a hydrographer. the expedition left tromsø on 30 june, visited jan mayen, and then made a difficult passage of the pack ice belt arriving at the east greenland coast on 25 july. working first in the tyrolerdal area, the veslekari went to the assistance of the polarbjørn which had run aground, then sailed south and west to the inner part of kejser franz joseph fjord, where work was carried out at the head of kjerulf fjord. rhedin fjord, alpefjord and narhvalsund were also visited. dif ficulties with the pack ice caused delays and diversions, but the veslekari came free of the ice on 25 august and set course for spitsbergen. the expedition’s results are fully described by boyd (1948). scientific results in east greenland included a general hydrographic chart of the region 72° to 74°n, as well as detailed hydrographic surveys of tyroler fjord, kjerulf fjord and narhvalsund. photo gram metic topographic maps were produced of parts of tyrolerdal and narhvalgletscher, as well as a planetable survey of agassiz dal. regional botanical studies were made, while geological work concentrated on aspects of glacial and quaternary geology. only a few new place names are associated with the expedition, mainly found in the geological re ports. 1937–38 søren richter’s expedition søren richter, an archaeologist who had twice overwintered with arktisk næringsdrift expeditions, led a three-man hunting group using the terrain south of kong oscar fjord. the expedition travelled up and back with the polarbjørn, except for peder sulebak who continued until 1939 hunting alone (giæver 1939; p.s. mikkelsen 1994). 1937–39 hermann andresen’s expedition hermann andresen and lars vemøy travelled up in 1937 with the polarbjørn to work the wollaston forland terrain. lars vemøy returned to norway in 1938, while andresen continued alone until 1939. the 1938–39 season was generally a poor trapping year for the norwegian hunters, but andresen had a record year with 642 foxes, the highest total ever re corded by a single trapper (giæver 1939; p.s. mik kelsen 1994). 1937–40 sigurd tolløfsen’s expedition in 1937 a six-man hunting expedition led by sigurd tolløfsen travelled up on the polarbjørn, but due to bad ice conditions could not reach their hunting terrain and returned home. four men went up in 1938, and occupied the hunting terrain between kuhn ø and dove bugt. three returned home in 1939, with eivind tolløfsen continuing alone from a base at jónsbu until 1940 (giæver 1939; p.s. mik kelsen 1994). 1938 louise a. boyd’s arctic expedition the 1938 expedition proved to be louise boyd’s last major expedition to east greenland. the veslekari, captained by johan olsen, was expedition ship, and scientists included a hydrographer, a surveyor and a geologist. leaving norway on 13 june, the veslekari 48 visited jan mayen on the way to the coast of east greenland which was reached at bass rock on 25 july. investigations were made around clavering ø and in granta fjord until 31 july, when the veslekari headed northwards along the coast. on 2 august the north-east end of île de france (now qeqertaq prins henrik; 77°48´n) was reached just south of kap montpensier (the belgica had reached 78°10´n in the pack ice in 1905, but their northernmost landing was on southern île de france). retreating southwards, parts of dove bugt were explored, and the inner parts of bessel fjord and ardencaple fjord visited. on 27 august the veslekari left the coast for spitsbergen. an account of the voyage is given by boyd (1948). the main scientific results included a general hydrographic chart of the region 74° to 77°n, with detailed profiles in pustervig and off soraner glet scher. tidal observations were made at danmark havn. other work included geological studies, botanical work and a survey of the orienteringsøer. 1938 sea-plane expedition to peary land: lauge koch supposed sightings of land between kronprins chri stian land and spitsbergen had been made by j.p. koch during the 1906–08 danmark-ekspedi tionen, by lauge koch in 1933 and peter freuchen in 1935. another alleged sighting of what had become known as fata morgana land by ivan d. papanin’s ice drift expedition in 1937 led directly to lauge koch’s 1938 seaplane expedition (koch 1940). koch flew to kings bay in spitsbergen with the dornier seaplane to be used on the two greenland flights, while the gustav holm sailed to kings bay with a reserve heinkel seaplane. the first flight on 10 may reached the coast of kronprins christian land, while the second on 15–16 may extended across peary land. both flights crossed the supposed position of the mysterious land sightings, but no trace of land was seen. 1938–39 ole klokset’s expedition this two-man norwegian hunting expedition, com prising ole klokset and a swedish assistent, was put on land by the sealer grande. a station was built on the north side of geographical society ø at kap mackenzie and huts built on the north side of ymer ø and east of walterhausen gletscher (pedersen 1969). 1938–39 den norsk–franske polar ekspedisjon (the norwegian–french polar expedition) willy knutsen and count gaston micard embarked on a combined hunting and scientific expedition in 1938. micard purchased the ringsel, which was renamed the en avant and captained for the voyage by karl nicolaisen. a main station, micardbu, and three huts were built on the east coast of germania land, and two huts on islands south of danmark havn. thirteen men overwintered, the en avant in winter harbour in northern lille koldewey. weather reports were sent to oslo three times a day. during the winter gaston micard became ill, and was evacuated by a stinson seaplane operating from the ship veslekari (knutsen 1949). 1938–39 den danske hundeslædeekspedition (the danish dog-sledge expedition): elmar drastrup in the winter of 1938–39, elmar drastrup and finn kristoffersen made a journey by dog sledge along the coast of east greenland from sandodden in young sund to ingolf fjord, and explored a new route to the interior of kronprins christian land. the purpose of the journey was to find a better land route to peary land, and if possible to traverse across to north-west greenland, although the latter objective was frustrated by open water and heavier than usual snow conditions that forced a retreat back along the east green land coast. a journey of 2350 km was completed in 105 travelling days. improvements were made to the map on the route of the expedition, especially in the interior of ingolf fjord and the valley system of vandredalen. sixteen place names, nine of them approved, are found in expedition reports (drastrup 1945; kristoffersen 1969). most names were given for incidents or the shapes of features, while the name vandredalen commemorates the probable migration route of musk oxen between north and east green land. 1938–39 mørkefjord expedition: eigil knuth & ebbe munck an alleged sighting of the mythical fata morgana land between spitsbergen and kronprins christian land by ivan d. papanin in 1937 was a prime factor in the promotion of this expedition, although its main aims came to be the exploration of the little known land region between latitudes 76° and 82°n, only 49 traversed previously by 1906–08 danmark-ekspedi tionen and the 1909–12 alabama expedition (knuth 1940, 1942). the somewhat cumbersome full name of the expedition led by eigil knuth and ebbe munck was ‘den danske nordøstgrønlands ekspedition, ud sendt af alf trolle, ebbe munck og eigil knuth til minde om danmark-ekspeditionen’ (the danish north-east greenland expedition, sent out by alf trolle, ebbe munck and eigil knuth to commemorate the dan mark expedition); the participants some times used an abbreviated form ‘munek-ekspeditionen’, but it is generally known as the ‘mørke fjord expedition’ after the main base at mørkefjord. alf trolle had made very substantial financial donations, while other support came from the carlsberg and tuborg foundations. ebbe munck and eigil knuth were coleaders of the expedition, knuth being in charge of the wintering party (five scientists and three greenl ander sledge drivers). the ship gamma was purchased, and captained by peder marcus pedersen departed from copenhagen on 19 june 1938 with a cargo including 70 dogs and a de havilland tiger moth aircraft fitted with floats. the coast of east greenland was reached near store koldewey, and the expedition and its equipment were unloaded west of hvalrosodden at the mouth of mørkefjord. the wintering house, mørkefjord station, was built here, while alwin pedersen, a zoologist loosely attached to the expedition, had his own small house at hvalrosodden. between october 1938 and march 1939 seven sledge journeys were made northwards to lay out depots for the spring sledge journeys, of which there were three between april and june. eigil nielsen reached the north point of kronprins christian land, exploring on the way the interior of ingolf fjord. eigil knuth reached as far as antarctic bugt, but also explored part of skærfjorden and the norske øer. svend sølver explored jøkelbugten, and penetrated westwards into the nunatak region climbing mile pælen on moltke nunatak. meanwhile, farther south, alwin pedersen and paul gelting made numerous shorter journeys around dove bugt, and to sælsøen and annekssøen. knuth (1942) lists 156 new place names, some with explanations of their origin. some of the fea tures named, especially around the mørkefjord station, are very minor. the great majority of the names are descriptive, given for the shape, colour or geographical position. about 15 commemorate per sons, including danish princes and princesses, and members of earlier expeditions. the main party returned home with the gamma in 1939, but mørkefjord station continued to be operated as a weather station by four men until 1942, although with increasing difficulty due to the war in europe. two men made a 1000 km journey from mørke fjord to scoresbysund in may–july 1940 (haar løv 1941, 1957). in april 1941 four men left the station to go south, leaving just ib poulsen and marius jensen. the last two men were evacuated by the northland in the summer of 1941. ib poulsen was to become leader of nordøstgrønlands slædepatrulje (the forerunner to the present sirius sledge patrol). 1939–40 swedish-norwegian expedition to east greenland this five-man expedition to clavering ø included the norwegian medical student kaare rodahl, who in vestigated vitamins in arctic diet, and the swedish professor hans w:son ahlmann, who carried out glaciological studies (rodahl 1943). three assistants, two of them norwegian hunters, accompanied the expedition. ahlmann and rodahl travelled up with the polarbjørn arriving in july 1939; ahlmann re turned with the ship in august 1939. rodahl re mained in east greenland until august 1940, when he went with theveslekari to iceland, and later to the orkney islands. the hunting station at revet was used as a base and laboratory, while a small hut was built in lerbugt on northern clavering ø. glaciological studies were carried out mainly on frejagletscher, and ascents were made of højnålen and moltke bjerg. rodahl’s biological studies led, amongst other things, to the discovery that poisoning due to eating polar bear liver arises from vitamin a enrichment. usage in the scientific publications of this expedition of several norwegian place names on northern clavering ø (lacmann 1937) led to their formal ap proval by danish authorities in 1950. 1939–40 søren richter’s expedition this three-man norwegian hunting expedition worked the terrain on the south side of kong oscar fjord. a new main station, havna, was built near noret and made the best catch of all the norwegian stations that winter, a total of 82 foxes, 34 of them kept alive in cages until their condition was optimal when they were killed. after the outbreak of war in 50 europe the hunters travelled to iceland in the summer of 1940. 1940–44 german meteorological expeditions when the danish and norwegian weather stations in east greenland ceased to transmit at the outbreak of war, germany attempted to establish its own meteorological stations in order to follow the development of weather conditions in the north atlantic. five main expeditions are recorded (holzapfel 1953) and are listed below, of which two operated radio stations for some time before being put out of action (ho warth 1957; olsen 1965). named features are associated with one of these, the 1943–44 operation bass geiger. 1940 the veslekari and furenak expeditions the first attempts by the german occupying powers in norway to obtain weather reports from east green land involved the sending of nazi sympathisers to east greenland with hunting personnel. bjerre (1980) records that the veslekari was sent to green land as usual to relieve the radio station at mygg bukta, but was arrested by the the fridtjof nansen, a norwegian patrol boat in allied service; the radio facilities at myggbukta were destroyed. the furenak was sent to east greenland from ålesund and landed a party of four danes on the south side of davy sund in the autumn of 1940; the party was discovered by the fridtjof nansen, while building a winter-house, and the house and installations were destroyed (lønø 1964; akre 1983; p.s. mikkelsen 1994, 2008). 1941 the buskø expedition the norwegian sealer buskø landed a small party of german meteorologists in peters bugt in the summer of 1941. the sledge patrol observed the buskø and alerted the united states coast guard ship north land which arrested the landing party. 1941–45 nordøstgrønlands slædepatrulje (north-east greenland sledge patrol) the first north-east greenland sledge patrol was formed in the summer of 1941 on the initiative of eske brun [1904–1987]. eske brun was then provincial governor (landsfoged) of north greenland, and when denmark was occupied he activated his emergency powers and moved to godthåb (nuuk) as head of a united greenland administration (see also be low). the sledge patrol was to consist of volunteers amongst the 27 danes and norwegians stranded in east greenland at the outbreak of the war (mainly hunters and staff at the weather stations), and initially comprised six danes, three norwegians and six greenlander dog drivers. their responsibility was to patrol the coast from 70° to 77°n and to prevent and report german activity. the sledge-patrol activities led to the discovery of the german meteorological expedition at hansa bugt in march 1943 (see ‘1942– 43 operation holzauge: the sachsen expedition’), as a consequence of which the patrol member eli knudsen was shot at sandodden, and the sledge patrol base at eskimonæs burnt down. a second ger man expedition at kap sussi on shannon was attacked by the sledge patrol in april 1944 (see ‘1943– 44 operation bassgeiger’). accounts of these events are given by malmquist (1955), howarth (1957), willoughby (1957), henry rudi (in: søren sen 1958), liversidge (1960), olsen (1965), bjerre (1980) and p.s. mikkelsen (1994, 2008). in 1943 a new patrol base was established at døde mandsbugten, replaced in 1944 by a larger station erected with usa assistance at sandodden. emer gency huts were built on maria ø and in blæsedalen. the sledge patrol was disbanded in 1945 but revived in august 1950, the forerunner of the present sirius sledge patrol. 1941–45 usa – northeast greenland task unit united states activities in the coastal waters of east greenland during the war years began with the agreement negotiated in 1941 by eske brun (head of the united greenland administration at nuuk / godt håb) and the danish ambassador in washington, henrik kaufmann, by which the usa agreed to pro tect greenland against foreign invasion. from 1941 three coastguard patrol boats (the northland, north star and bear) were on duty in east green land under the command of edward h. smith (‘ice berg smith’), and to some extent supported and sup plied nordøst grønlands slædepatrulje (willough by 1957). in 1944 the patrol boats were partly replaced by the icebreakers eastwind and southwind. in 1944 the northland sank the kehdingen, and the two icebreakers captured the externsteine; both ships had been carrying german meteorological expeditions. 51 1942–43 ‘operation holzauge’: the sachsen expedition a 19-man german meteorological expedition transported aboard the sachsen landed in hansa bugt in august 1942, and operated undetected until march 1943 when members of the sledge patrol met a group of german soldiers. in subsequent encounters, eli knudsen was killed at sandodden, eskimonæs sta tion was burnt down, and the leader of the german party, lieutenant herman ritter, was captured and taken to scoresbysund (howarth 1957). the hansa bugt weather station was bombed by four b-24 air craft on 25 may 1943 (balchen 1958), causing some damage, and leading to evacuation of the per sonnel by a german flying-boat between 7 and 17 june. the ship sachsen was burnt, and other installations de stroyed. one member of the german expedition accidently left behind (rudolf sensse) was taken prisoner by the northland in july. an account of events from the german side is given by weiss (1949). a different interpretation of events is given by bjarne akre (1983). the two norwegian akre cousins in the sledge patrol were unhappy with their danish colleagues, and disagreed with just about every decision that was made by ib poulsen, the sledge patrol leader. the account by akre suggests that eli knud sen and ib poulsen were actually nazi-sympathisers, and that lieutenant herman ritter (the german com mander captured and taken as a prisoner to scores bysund) may have been an imposter, perhaps the norwegian captain sverre strøm whom they had met in ivigtut the previous year. this strange story does not seem to have aroused much interest in denmark, and bjarne akre does not make his case more plausible by constantly referring to the leader of the sledge patrol as ‘palle’. 1943–44 operation bassgeiger a german meteorological expedition of 27 men aboard their ship coburg was frozen in off kap sussi on the outer coast of shannon in october 1943. the coburg was eventually crushed by the ice and abandoned. the expedition established a subsurface base camp in a snow fan at kap sussi, which on 22 april 1944 was attacked by members of nordøstgrønlands slædepatrulje. the only casualty was gerhard za c her, a german lieutenant, who was buried at kap sussi (fig. 16). the expedition was evacuated by ger man flying-boat on 3 june 1944. olsen (1965) de scribes the events, and also notes nine place names used by the expedition for localities in the im mediate vicinity of the base camp. german accounts of this operation include those of triloff (1948) and schatz (1951). a recent detailed and well-illustrated ac count (in danish) is provided by frederiksen (2008). 52 fig. 16. on 22 april 1944 the nordøstgrønlands slædepatrulje (sledge patrol) attacked the german meteorological station established in a large snow drift near kap sussi on shannon. the only casualty was gerhard zacher, a german lieutenant, whose grave lies undisturbed at kap sussi. following the attack the german expedition was evacuated by air on 3 june 1944. 1944 operation edelweiss an attempt was made by the kehdingen to land a german meteorological expedition in 1944, but it was intercepted by the us patrol boat northland near the south point of store koldewey, and sunk (wil loughby 1957; liversidge 1960; olsen 1965). the crew of 28 was taken prisoner. 1944 goldschmied expedition: operation edelweiss ii this 12-man german meteorological expedition reached land on the east side of lille koldewey on 1 october 1944 (fig. 17). the landing party was cap tured on 4 october by troops from the us icebreaker eastwind. the expedition ship ex tern steine was trapped in the ice and subsequently captured by the eastwind and southwind; it was unofficially re named the eastbreeze (willoughby 1957; liversidge 1960; olsen 1965), and later became uss callo. 1945–52 østgrønlandsk fangstkompagni nanok a/s (east greenland trapping company nanok ltd.) the danish hunting company nanok resumed hunting activities in 1945. their huts were then in a poor state of repair after the ravages and neglect of the war years, although the danish government did pay compensation for the use of the huts and provisions during the war years, and continued to pay an annual subsidy until 1951. between 1945 and 1951 a total of 23 huts were built, as well as new stations at the head of loch fyne and at germaniahavn. the suspension of subsidies was related to the establishment of slæde patruljen sirius in 1950 which was henceforth to be the official danish presence in east greenland. by the summer of 1952 only one danish hunter re mained in east greenland, and 1952 effectively marked the end of danish trapping. j.g. jennov had visited east greenland virtually every summer since the war, and his last visit was after the end of hunting, in 1954, when he rescued mønstedhus from falling into the sea; it was moved 20 m to safety. brief ac counts of post-war activities are given by lauritsen (1984) and p.s. mikkelsen (1994, 2008). 1946–59 arktisk næringsdrift a/s (arctic commercial enterprise ltd.) arktisk næringsdrift resumed hunting operations in 1946, with the aid of a norwegian state subsidy towards hire of the annual relief ship, and an interest-free loan. many hunting huts and stations were in poor condition, partly due to neglect and partly due to deliberate destruction during the war years. mygg bukta weather station was repaired and weather re ports resumed in august 1946. in 1948, a replacement for the destroyed jónsbu radio station was built. however, the northern stations of ottostrand and ny jónsbu were given up in 1953, due to poor hunting and difficulties of access. in 1959 the norwegian state suspended its subsidy to the weather station at mygg bukta, and this, together with falling skin prices and the increasing cost of ship hire led to a cessation of norwegian hunting. the polarbjørn was the relief ship from 1946 to 53 fig. 17. depot of fuel drums laid out by a german meteorological expedition at røseløbet, lille koldewey, on 1 october 1944. the 12-man landing party was captured on 4 october by troops from the us icebreaker eastwind. 1948, the quest in 1949, and the new polarbjørn from 1950 to 1957. in 1957, the polarbjørn was crush ed in the ice and lost, the crew and passengers being rescued by the danish naval cutter teisten and flown home from mestersvig. in addition to the hunters, the norwegian ships occasionally transported scientific and climbing expeditions to east green land, and in the later years a few tourists. fox hunting was very poor in 1948–1949 and 1955–1956, and catastrophic in 1956–1957 when hunters at myggbukta, hoelsbo and revet had to gether a catch of only 36 foxes. salmon fishing was undertaken in some years, sometimes with success, sometimes with disastrous results. lønø (1964) de scribes the post-war norwegian hunting activities, and reported the total catch of arktisk næringsdrift from 1946 to 1959 as more than 5000 foxes and 40 bears. the danish –norwegian agreement on east green land was terminated in 1967, and in 1969 the danish state took over the 150 norwegian hunting huts and stations paying danish kroner (dkk) 50 000 in com pensation. 1946–59 hermann andresen’s expeditions hermann andresen, a norwegian hunter who had last overwintered in 1938–39, organised a series of expeditions to what norwegian trappers called the ‘sunnmøring terrain’ from 1946 onwards. kap her schell was the main station in the north, and in the south the stations at antarctic havn, havna and kap peterséns were used. andresen received a state subsidy in 1946 to repair the old huts and build new, and received further annual subsidies subsequently. three or four hunters were active each year, altogether 32 men with a total of 42 winters between them. four hunters in the southern terrain broke their contracts in poor hunting seasons, taking work at the lead mine near mestersvig. from 1948 andresen also organised summer salmon (arctic char) fishing, sending up to five men with the relief ships to fish, mainly in the rivers at brogetdal, zackenberg, dusén fjord and loch fyne. together with arktisk nærings drift, 358 barrels of salmon were taken between 1937 and 1959 (lønø 1964). andresen’s expeditions were dependent on arktisk næringsdrift for transport to and from greenland, and were also obliged to suspend activities in 1959. 1947 united states air force photo grammetric flights photogrammetric flights were made in 1947 over east greenland, as well as the greater part of the ice-free areas of other parts of greenland, by the united states air force. the oblique and vertical aerial photographs obtained were used to produce the 1:250 000 scale map sheets of the army map service (ams), the east greenland sheets being compiled in 1952. the usaf aeronautical charts at the same scale used the same database, but with altitudes and contours in feet rather than metres. 1947–50 dansk peary land ekspedition (danish peary land expedition) the main area of activity of this expedition, one of the series of expeditions to peary land led by eigil knuth, lies in north greenland, north of the area of interest of this volume (martens et al. 2003). how ever, a southern base of the expedition was established at zackenberg bugt in young sund (74°28´n); every year equipment and expedition members were sailed to the base by the godthaab. catalina sea planes were used to ferry stores and personnel to peary land. eigil knuth frequently used the incorrect one word name ‘pearyland’ when referring to the activities of his expeditions. opportunity was taken by some expedition mem bers to carry out archaeological and other work around the southern base. in addition other expeditions took advantage of the transport possibilities of the godthaab to reach east greenland. the latter included the 1948 leeds university greenland expedition, the 1949 w.r.b. battle expedition and the 1951 british north greenland (reconnaissance) expedition. 1947–58 de danske ekspeditioner til østgrønland (the danish expeditions to east greenland): lauge koch lauge koch’s expeditions to east greenland resumed in 1947, with government support and on a more re gular basis than pre-war, and with an almost entirely geological bias (koch 1961). their format was at first similar to the last pre-war expeditions, based on ships with groups of scientists overwintering. how ever, catalina flying boats soon replaced ships for transport of personnel, and after 1952 when the air port 54 was constructed at mestersvig, dc-4 aircraft were used. in 1948 the expedition acquired its first norse man seaplane, and in 1949 a second norseman (fig. 18). overwintering was given up in 1953. koch re cords that 691 persons took part in his post-war expeditions, but this figure included in addition to scientists, the crews of the boats, and the mining engineers and drilling teams involved in prospecting around mestersvig. in general six to eleven geological teams were active each year. compilation of geological maps was begun in 1955 by john haller, and to complete these maps and fill out gaps, more than 32 000 km of reconnaissance and photographic flying was carried out with the two norseman aircraft in 1955, 1956 and 1958. john haller’s compilation work continued after the expeditions stopped in 1958, and the geological maps – printed in 1964 – were pub lished in 1971 (koch & hal ler 1971). a major geological account of the east greenland caledonides was published the same year (haller 1971). a brief sum mary of each year’s activities is given below. 1947 – the expedition was based on the gustav holm, and comprised 30 members including four geological parties; it was active between latitudes 72° and 74°n. 1948 – gustav holm and one norseman seaplane provided transport for 47 members including eight geological parties. the area of activity was again from 72° to 74°n, and lead and zinc deposits were found near mestersvig. 1949 – the expedition comprised 97 members including seven geological parties, and was supported by the gustav holm with two seaplanes for transport and reconnaissance. icelandic ponies were used for the last time. special attention was given to the lead mineralisation near mestersvig. 1950 – catalina and norseman aircraft were used to transport the 120 members of the expedition, which included nine geological parties and 86 pros pecting and drilling personnel. the ships gustav holm, veslekari and polarstjerne were used to transport equipment and materials for the prospecting group. erdhardt fränkl made one of the earliest explorations of the stauning alper, and gerold styger made ascents in the werner bjerge. 1951 – catalina and norseman aircraft were used to transport the group of 104 to east greenland, the numbers including 58 prospecting and drilling per sonnel. eight geological parties were active between 70° and 74°n. one party, including eduard wenk and john haller, climbed petermann bjerg and other near by peaks during geological mapping (wenk & haller 1953; buess 1953), and a second party led by hans r. katz made a journey to the nunatak region at 74°n in ‘weasel’ tractors of paul-emile victor’s expedition (see ‘1950–51 expéditions polaires françaises, missions paul-émile victor’), supported by an air drop at cecilia nunatak (fig. 19; katz 1951; diehl 1953). fränkl continued his explorations in the northern stauning alper, making first ascents of fri heds tinde and elisabethsminde. 1952 – the expedition numbered 49, including eight geological parties, and was transported by cata lina and norseman aircraft. two parties worked from a base at centrumsø in kronprins christian land (80°10´n). a two-man group overwintered at ella ø from 1952 to 1953, after which wintering was given up (fig. 40). west of the bay known as mesters vig an airfield was constructed (subsequently known in the one-word form mestersvig), and the newly formed mining company, nordisk mine selskab, be gan ex ploi tation of the lead deposits. 55 fig. 18. norseman aircraft of lauge koch’s expedition that was used extensively in the 1950s for aerial photography and geological reconnaissance flights. the john haller photograph collection, geus archive. 1953 – catalina and norseman aircraft transported 41 expedition members to greenland, including seven geological parties. two parties again worked out of centrumsø, one of them flying northwards to peary land by catalina, and traversing the north greenland fold belt to reach kap morris jesup (the northernmost point of the greenland mainland). another party made a long journey to the nunataks west of goodenough land, including an ascent of shackleton bjerg (haller 1954), and south of mesters vig molybdenum was discovered at malmbjerg. 1954 – catalina, norseman and dc4 aircraft trans ported 39 personnel to greenland, including nine geological parties (christensen 1955). one par ty, including john haller, wolfgang diehl and fritz schwarzenbach worked in the stauning alper and made several major ascents, including danske tinden and norsketinden (diehl 1956). 1955 – catalina and norseman aircraft transported 34 members to east greenland. there were seven geological parties working over a wide area between 70° and 78°n. two parties supported by norseman aircraft worked out of a base at krumme langsø (75°03´n). extensive reconnaissance and photographic flights were made with norseman aircraft out of satellite bases at daneborg, krumme langsø, dan mark havn and britannia sø. 1956 – catalina, norseman and dc-4 aircraft were used to transport the 33 personnel to greenland. these included eight geological parties, two of which worked between 70° and 72°n. two sikorsky helicopters were used in co-operation with nordisk mine selskab, and extensive aerial reconnaissance and pho to graphy were carried out with norseman aircraft between bessel fjord and the stauning alper. 1957 – norseman, catalina and dc-4 aircraft trans ported 47 expedition members to east green land. five of the 11 geological parties worked between 70° and 72°n. 1958 – catalina, norseman and dc-4 aircraft transported 55 members of lauge koch’s last expedition to east greenland. eight of the 11 geological parties worked south of 72°n. some extended reconnaissance and photographic flights were made. expeditions had been planned to complete the mapping of the scoresby sund region (70° to 72°n) from a base at rypefjord, but financing of lauge koch’s expeditions was unexpectedly brought to an end after the 1958 season. lauge koch’s post-war expeditions were responsible for the introduction of about 550 new place names in east greenland. the minutes of the place name committee for this period are almost complete, and nearly all the names can be attributed to specific geologists. however, the origin of the names is not always apparent. peter bearth worked in the werner bjerge region in 1953 and 1954, and gave about 70 names (bearth 1959). many were given for the shape and character of features, some for events, some with geological con nec tions, while a few commemorate swiss geologists. heinrich bütler took part in expeditions in 1948, 1950 and from 1952 to 1957, but appears to have been directly responsible for only two new names. john w. cowie took part in expeditions from 1949 to 1954, and gave five names in the ella ø region, 56 fig. 19. in 1951, h.r. katz and two companions were transported by weasel snow tractors of paul emile victor’s expedition from cecilia nunatak westwards and northwards into the nunatak region at about 74°n. after they were dropped off, valuable geological observations were made while making their way eastwards on skis. at eleonore sø, at about 1600 m above sea level, the party abandoned their skis due to lack of snow on the glaciers. this depot was found by a ggu expedition in 1975. most of them commemorating persons (cowie & adams 1957). desmond t. donovan worked mainly on traill ø during five summers between 1947 and 1957. he is credited with 25 names, some with geological con nec tions, two for the english towns of bath and bristol, and two for noted british geologists (donovan 1964). silvio eha took part in expeditions from 1947 to 1949, working mainly on ymer ø and in lyell land. his 15 place names were mainly given for the shape or character of features, or for events during the expedition (eha 1953). erdhardt fränkl gave 53 names to features following his work between 1948 and 1953. of these 11 were at about 80°n, while others were in the stauning alper and andrée land regions (fränkl 1953, 1954). most were given for characteristics of the features or events during the expeditions. p. graeter gave four names to various features on gauss halvø following his work in 1950. john haller worked throughout northern east greenland in the years 1949–56 and 1958, and made many geological reconnaissance and photographic flights together with ernst hofer (fig. 20), using norseman aircraft. he gave 154 names in the region 71°–79°n, including some commemorating scottish castles, some for austrian geologists, some for mem bers of the 1906–08 danmark-ekspeditionen, a group with geological connections, a few for swiss mountains and a few for the shapes of features (haller, 1953a, b, 1955, 1956, 1958; wenk & haller 1953). m.y. hassan worked up collections made at kap brewster for f.w. sherrell, and used four new names, three with geological connections (hassan 1953). hans p. heres worked on south-east traill ø in 1956–58, and his 16 names were given mainly for the shape or character of features, with one commemorating countess maria-theresia of austria. hans kapp took part in expeditions from 1955 to 1958, and gave 27 names to features in northern scores by land (kapp 1960). most were given for the shape or character of features, or for geological con nections, with a few commemorating incidents during the expedition. hans r. katz mapped areas in hobbs land and strindberg land in the years 1948, 1949 and 1951, giving 44 names (katz 1952). most relate to the shape of features, to geological characteristics or events during the expeditions. enrico kempter took part in the 1956 to 1958 expeditions, and gave 16 names to features north of sydkap, mainly for natural characteristics of the fea tures and their geology (kempter 1961). david malmquist gave four names to features at c. 79°n, including the eli knudsen øer. paul stern took part in the 1955 to 1958 expeditions, and is credited with five names. peter vogt mapped parts of hinks land in the years 1956–58, and gave four names, one of them for peter freuchen (vogt 1965). eduard wenk took part in expeditions in the years 1951 to 1954 and 1957 to 1958. he was responsible for 26 names, a number given for their appearance, a group with greek connections originating from his greek assistant, while a few have swiss origin (wenk 1961). hans zweifel mapped nathorst land in 1954 and 57 fig. 20. ernst hofer, a photographer from the swiss topographical institute, was employed by lauge koch’s geological expeditions to east greenland from 1951–54. hofer accompanied john haller on many reconnaissance flights with the expedition’s norseman aircraft, and his photographs illustrate his book “arctic riviera” (hofer 1957). 1955, and proposed 21 names (zweifel 1958). some record natural characteristics of the features, while a few have swiss connections. the prospecting activity near mestersvig in the years 1949 to 1851 led to preparation of 1:50 000 scale topographic maps, and the introduction of 48 place names. most of these relate to the prospecting and lead mineralisation, while some record the shape of features, and a few commemorate danish personalities. [place name committee archive.] 1948 leeds university greenland expedition: w.r.b. (ben) battle a four-man expedition from leeds university led by w.r.b. (ben) battle travelled to east greenland with the danish peary land expedition (see above, ‘1947– 50 dansk peary land ekspedition’) aboard the godt haab, arriving at zackenberg bugt at the end of july. a base camp was established in tyrolerdal west of the head of tyrolerfjord, where the expedition divided into two parties. one group undertook glacio l ogical studies on pasterz and nearby glaciers (battle 1952), while the second group made a general geological reconnaissance extending north to grand jean fjord (leedal 1952). fourteen new names were proposed for valleys, mountains and glaciers in the valley system of northern payer land and a.p. olsen land. the names were given mainly for natural features, while a few com memorate leeds university and cambridge colleges (e.g. ledesia bjerg, trinity gletscher). 1948–present danmarkshavn icao weather station the international civil aviation organization (icao) weather station danmarkshavn (spelt in one word) was established in 1948 at danmark havn. the ship g.c. amdrup had transported 400 tons of materials and 17 carpenters and radio personnel to the site during the summer, and the veslekari brought in more equipment in august the same year. the first station leader was ib poulsen (leader of the sledge patrol during the war years). operations were begun in the autumn of 1948, and the station was to have been completed in 1949; however, the g.c. amdrup was diverted due to difficult ice conditions to nord fjord, and exchange of personnel was carried out using norseman and catalina aircraft (thomsen 1966). the weather station is normally supplied in alternate years by ship from denmark, although it is not unusual for ice to prevent access. the staff at the station has given a number of place names to local features 1949 w.r.b. (ben) battle expedition originally planned as a four-man cambridge univer sity expedition, battle travelled up alone due to lack of space aboard the godthaab, the expedition ship of the dansk peary land ekspedition (see above, ‘1947–50 dansk peary land ekspedition’). glacio logical work was carried out on several of the glaciers on clavering ø (battle 1952), and the work led to formal approval of four norwegian names for clave ring ø glaciers (lacmann 1937), all derived from norse mythology. 1949–54 geodætisk institut (geodetic institute) aerial photography and surveying in 1949 low-level, vertical, aerial photography was car ried out in the region around mestersvig for the danish geodetic institute, with the main purpose of constructing detailed topographic maps in connection with the lead-zinc prospecting. oblique aerial photography was also carried out over much of the region between 69° and 81°n be tween 1950 and 1952. in 1951 a geodetic institue surveying party based on the ole rømer visited the scoresby sund region. this project continued in 1953 and 1954, with larger parties based on the tycho brahe and with helicopter support. in 1953 a helicopter technician was killed in an accident, a tragic incident commemorated by the name c. hofmann halvø. 1950–51 plankton studies in scoresby sund: peter digby peter s.b. digby and his wife vi, who had travelled to scoresbysund (70°29´n) with the jopeter in august 1950, made regular plankton hauls in the waters of scoresby sund between august 1950 and august 1951 from a small boat and through holes in the ice (digby & digby 1954). they lived in lauge koch’s ‘expedition house’ at scoresbysund built in 1926. digby re turned home with the jopeter in august 1951, his wife having flown home in july with their newly born baby. 58 1950–51 expéditions polaires françaises, missions paul-émile victor (french polar expeditions) paul-émile victor embarked in 1948 on a long series of expeditions to investigate the inland ice of green land, including meteorological, geophysical and glaciological observations. seismic and gravity sur veys were made over an extensive region between 63° and 74°n (fristrup 1966). in 1950 victor’s ‘weasels’ (powerful snow tractors) reached cecilia nunatak (72°30´n) in east greenland, and some of the expedition members made their way to ella ø and re turned to europe with lauge koch’s expedition. in the summer of 1951 a group of lauge koch’s geologists, led by hans r. katz, was transported by victor’s ‘weasel’ tractors from cecilia nunatak to the nunatak region near hobbs land at 74°n. katz and his party undertook a strenuous tour by ski and on foot eastwards to the coast of nordfjord. 1950–present: slædepatruljen sirius (sirius sledge patrol) the sledge patrol, which had operated in east green land during the war years, was re-established in august 1950. this followed the realisation by nato (north atlantic treaty organization) of the strategic significance of northern east greenland in the event of war, and some concern as to whether denmark was doing enough to uphold its rights of sovereignty over the unoccupied regions of north and east greenland. the patrol was known at first as ‘operation resolut’, and had a base at ella ø. in l951 it changed its name to ‘slædepatruljen resolut’, and moved to new headquarters at daneborg. a last name change to ‘slæde patruljen sirius’ (sirius sledge patrol), in common parlance ‘sirius’, was made in 1953, the name being given after the brightest star in the constellation canis major. sirius is a danish military police force which patrols the uninhabited regions of north and east greenland, roughly corresponding to the boundaries of the present day north-east greenland national park (nordøstgrønlands nationalpark). during the winter and spring dog-sledge teams cover a total of 20 000 km on patrol. occasional use is made of the old danish and norwegian hunting stations, but these have largely been replaced by prefabricated bear-proof huts. during the short summers, depots are laid out by aircraft and boat, and damaged huts repaired. widespread damage to the old hunting huts by bears in search of food means that few huts now survive in their original form. small military groups maintain the airfields at station nord and mesters vig. recent accounts of the activities of the patrol are given by bjerre (1980) and p.s. mikkelsen (1986, 2005). 1951 norwegian climbing expedition a party of three norwegians, a.r. heen, k. barstad and ø. roed, climbed three peaks in the northern stauning alper from a base at kap peterséns (72°25´n). these were the first ascents of tårnfjeld and varde fjeld, and the second ascent of elisabeths minde (ben net 1972). 1951 british north greenland expedition – reconnaissance: c.j.w. (james) simpson as a guest of the dansk peary land expedition in 1950, c.j.w. simpson had observed from a distance the largely unexplored nunataks of dronning louise land (76°–77°15´n), and considered the region as a suitable goal for a major british joint services expedition. a reconnaissance expedition in 1951 to check its possibilities was led by simpson. in july a depot was air-dropped on dronning louise land and a fourman group was landed by sunderland flying boat on sælsøen. accompanied by a trapper from hvalros odden (orla jensen), a journey was made across stor strømmen to dronning louise land where a site for a base was found on the shores of britannia sø. after limited exploration, the party recrossed stor strøm men and was picked up from sælsøen at the end of august (simpson 1955, 1957). 1952–54 british north greenland expedition: c.j.w. (james) simpson this major expedition to dronning louise land (76°–77°15´n) led by commander c.j.w. simpson was a co-operative venture involving all three branches of the british armed forces, the shell petroleum company and civilian scientists. the name of the expedition is a misnomer, as dronning louise land is a long distance from ‘north greenland’. the expedition in the field numbered 30, eight of whom returned home in the summer of 1953, while an additional five members took part only in the second year. the objects of the expedition included a comprehensive scientific programme, as well as providing mem 59 60 fig. 21. the british north greenland expedition established a base at britannia sø, dronning louise land, in 1952. the expedition was supplied by air, and carried out scientific investigations throughout dronning louise land until the summer of 1954. this simplified map shows the new place names given by the expedition as well as earlier names (from: peacock 1958). bers of the armed forces with arctic experience. glaciological, meteorological, physiological and geo physical studies were carried out (fig. 21). the meteorological work included establishment of a station, ‘northice’, at the centre of the inland ice west of dron ning louise land, while the geophysical work involved a traverse from dronning louise land across the inland ice to thule in north-west green land. accounts of the expedition include those of sim pson (1955, 1957), banks (1957) and hamilton (1958). the british armed forces provided air transport, equipment and many of the expedition mem bers, while financial backing came chiefly from the shell petroleum company and a personal contribution from sir winston churchill. in july 1952 the norwegian sealer tottan sailed equipment to the southern base at zackenberg bugt in young sund, on its first journey sailing via ivittuut (ivigtut) in west greenland to pick up dogs. in early august most of the expedition members and their equipment were air-lifted to britannia sø by sunder land aircraft, and a main base was established on the north shore of the lake. eight ‘weasel’ snow tractors, too bulky to be carried by air, were landed at kap rink (75°08´n) by the tottan in late august. while waiting for the ice to freeze, several peaks were climbed in the nearby barth bjerge. with assistance from the danish personnel at danmarkshavn and members of sirius, the group with the ‘weasel’ tractors made the journey to danmarkshavn in the au tumn. meanwhile ‘northice’ had been established with the aid of airdrops from thule, in the course of which a hastings aircraft crash-landed. while surveying in april 1953, the danish mem ber of the expedition, hans a. jensen, was killed in a fall near kap niels (76°23´n). the eight ‘weasel’ tractors made a difficult journey to britannia sø via sæl søen and storstrømmen, and in may began their jour neys on the inland ice. new supplies were brought into young sund (74°27´n) in early august 1953 by the polar sirkel, and air-lifted to britannia sø together with the five new expedition members replacing those leaving. surveying and geological exploration was carried out on numerous journeys throughout dronning louise land in 1953 and the first half of 1954. in august 1954 the entire expedition was evacuated from britannia sø, apart from members of the gravity team who returned home from thule, after their crossing of the inland ice. sixty new place names were proposed as a result of the expedition (fig. 21), given mainly for notable phy sicists, musical composers, and organisations or individuals who had given substantial assistance to the expedition. an additional 12 unapproved variations of names occur in expedition reports. 1952–90 nordisk mineselskab (northern mining company) following discovery of lead and zinc mineralisation in the mesters vig region (72°13´n) by geologists of lauge koch’s expeditions in 1948, the northern mi n ing company (nordisk mineselskab) was established in 1952; it was commonly known by the abbreviated name ‘nordmine’. originally 27.5% of the company was owned by the danish state, the balance being held by danish, swedish and canadian interests. an exclusive concession covering the region 70° to 74°30´n was granted in 1952 for a period of fifty years. detailed studies of the lead-zinc showings were commenced in 1952, and in the following years a mining town was built in blydal, a road built between the town and the harbour (nyhavn), and underground workings opened. production commenced in 1956 and the mine was worked out by 1962. ap proximately 545 000 tons of concentrate (9.3% pb and 9.9% zn) were shipped out, with expenses roughly balancing earnings (thomassen 2005a). the airfield known as mestersvig, which was opened in 1952 to serve the mine, remained open for general use until 1985 – when it was replaced for most purposes by a new airfield built at constable pynt (70°44´n). in 1958 diamond drilling was commenced at a new prospect known as malmbjerg (71°59´n), where lauge koch’s geologists had reported molybdenum mineralisation in 1954. further drilling was carried out in 1959 and 1960, after which the company ark tisk mineselskab was formed to continue investigations. a concession to exploit molybdenum and re lated minerals was granted in 1961, originally for a period of fifty years, but following extensive negotiations was relinquished in 1984. from 1968 to 1972 extensive regional prospecting was carried out throughout the nordisk mineselskab concession area (70°–74°30'n), in many years with helicopter support. preliminary oil exploration studies were carried out in 1971 and 1972 in cooperation with the atlantic richfield company (arco), but these were suspended as they appeared to be in breach of the terms of the original concession. regional prospecting activities were continued 61 from 1974 to 1976 and 1979 to 1984. from 1979 to 1982 investigations had financial support from the eec (european economic community), and led to finds of widespread scheelite. another eec-supported project in 1983 and 1984, to study tungsten-antimony mineralisation on ymer ø, included drilling at two localities in margerie dal (73°09´n). extensive negotiations in 1983–84 concerning concessions to explore for and exploit oil and gas in the jameson land basin (70°30´–72°n) by nordisk mineselskab and arco led to granting of an exclusive concession in 1984 (see below). at the same time the original nordisk mineselskab concession rights were relinquished, and replaced by six exclusive mine ral concessions and one concession for hydrocarbons. however, these concessions lapsed when nordisk mineselskab closed down in 1990. 1953–64 grønlands zoogeografiske undersøgelse (zoogeographical investigations in greenland): christian vibe in 1948 christian vibe was appointed head of grøn lands zoogeografiske undersøgelse and was based at the zoological museum in copenhagen. his travels to greenland were directly funded by the ministry for greenland, and between 1953 and 1964 he made six visits to east greenland (vibe 1967). 1953 – christian vibe visited the region around mestersvig (72°13´n) to study birds and mammals. 1954 – vibe returned to east greenland with the specific objective of capturing musk-ox calves that were to be transferred to west greenland. however, reconnaissance flights in jameson land, andrée land and ymer ø (71°–73°n) revealed very few calves and a very high death rate among musk oxen in the winter of 1953–54. 1956 – christian vibe visited the scoresby sund region to study the population of musk oxen. jameson land and liverpool land were traversed on foot, and the interior branches of the fjord system were overflown using catalina aircraft. 1958 – christian vibe and torben andersen visited the scoresby sund region (70°–72°n) to continue studies of musk oxen. in co-operation with lauge koch’s expedition, large areas were overflown by catalina, with landings in gåseland, charcot land and rypefjord (andersen 1960). 1961 – christian vibe again visited the scoresby sund region (70°–72°n) with six assistants and with the purpose of capturing musk-ox calves. twelve calves were captured at rypefjord, and a further two at daneborg. they were taken back to copenhagen aboard the kista dan. one died soon after arrival, and in 1962 the surviving 13 calves were transferred to the søndre strømfjord region of west greenland (nielsen & küter 2000). 1964 – christian vibe assisted by j. van hauen, j. højsgaard, c.c. scavenius and others, captured 16 musk-ox calves and two yearlings in rypefjord (71°n). these were sailed to copenhagen with the thala dan. two calves died and two others were sick in early 1965, but the remaining 14 were transferred to søndre strømfjord to join the group sent there in 1962. this original 27-strong group of musk oxen bred so successfully that its numbers had risen to 200 by 1980, 1000 in 1985, and the population was estimated at 4000 in 1999 (nielsen & küter 2000). in 1986 a number of yearlings from the søndre strøm fjord population were flown to inglefield land, with the intention of forming a new breeding group. 1954 danish–norwegian expedition to the stauning alper a four-man climbing expedition explored the vi king e bræ region of the stauning alper (72°n). three participants (a.r. heen, ø. roed and e. jensen) took part in the ascent of their main objective, norske tinde, which they originally called eirik rødes tinde or stortoppen (hoff 1955; bennet 1972). two lesser peaks overlooking alpefjord (hellefjeld and skifer bjerg) were also climbed. 1955 cambridge expedition to east greenland: j.b. latter an ornithological expedition of five led by j.b. latter visited antarctic havn and fleming fjord (71°40´– 72°n) in late july and early august, and succeeded in ringing 11 pink-footed and 299 barnacle geese (latter 1956). three members of the party were from cam bridge university (uk), and one each from oslo uni versity (norway) and birmingham university (uk). 1955 geodætisk institut (geodetic institute) name registration a party of two from the danish geodætisk institut (captain j. balle and e. laursen) were sent to scores bysund / illoqqortoormiut (ittoqqortormiit) (70°29´n) in 1955 to record place names used by the greenlandic 62 population in the region, a procedure also carried out by the geodætisk institut in other parts of green land. approximately 190 names were registered, nearly all of them of the typically descriptive type, some of which clearly originated from the earliest days of the settlement and were still in use. a further 10–15 names have been introduced in modern times, reflecting the changing use of the resident green landers. the east greenland dialect dif fers from that of west greenland, and differences are sometimes reflected in the place names. names are listed in this volume according to the new orthography (spelling reform) that came into use in 1972, but cross-references from the old spelling still found on many published maps are included. [place name com mit tee archive.] 1955–64 mestersvig geomorphological research: a.l. (linc) washburn a.l. washburn embarked in 1955 on a long-term programme of geomorphological studies from a base adjacent to the airport at mestersvig (72°13´n), in association with h.m. raup, f. ugolini and other scientists at different times. reconnaissance studies in 1955 and 1956 were followed by the main phase of the study which lasted from 1957 to 1961, with fol low-up studies in 1964 (washburn 1965). the head quarters of the expedition was at camp tahoe, a house north of tunnelelv on a section of road be tween nyhavn and minebyen; this house has subsequently become known as washburn’s hus. 1956 cambridge university and marlborough college expedition to north-east greenland: g. thomas wright a party of six led by g. thomas wright visited the hold with hope area (74°n) in late summer to make observations of pink-footed and barnacle geese. the party travelled up with the polarbjørn, and visited many localities between loch fyne and ymer ø. however, the cold spring and early summer meant that 1956 was a non-breeding year for geese, and only four were ringed. the party was picked up by the polarbjørn, that was escorting the salvaged sealer jopeter back to europe (wright 1957; goodhart & wright 1958). 1956 mountaineering in the werner bjerge: w.d. brooker a party led by w.d. brooker is reported to have climbed two peaks in the werner bjerge (including malmbjerg; 72°n), and two peaks in the stauning alper 1523 m and 1676 m high (fantin 1969, p. 71). 1956 ‘operation defrost’: s.m. needleman a four-man party led by s.m. needleman carried out a reconnaissance survey of north greenland for the air force cambridge research center to locate potential aircraft landing sites. the centrumsø region (79°n) of kronprins christian land was visited from 15–18 august (needleman 1962). investigations were continued in 1960 as ‘operation groundhog’. 1957 austrian east greenland expedition (die österreichische grönlandexpedition): hans gsellman a party of eight austrians, led by hans gsellman, visited furesø and the stauning alper region (72°n). the party flew by catalina directly to the dammen region of inner alpefjord. two men made a boat trip to the west end of furesø and climbed a peak overlooking violin gletscher. sefström gletscher was explored and a total of 19 summits were climbed, including 11 first ascents. the latter included sef ström tinde and sefströmsgipfel, both over 2700 m high. their nine other first ascents were named mainly for their appearance. the party had difficulty leaving the area, and were eventually transported from dammen to mestersvig aboard the small boat the netta dan (also referred to as ‘vippa dan’) owned by the danish ship-owner knud lauritsen (gsellman 1958a, b; koglbauer 1965; bennet 1972). 1958 scottish east greenland expedition: c.m.g. (malcolm) slesser a nine-member climbing expedition led by c.g.m. (malcolm) slesser explored the bersærkerbræ and sefström gletscher area of the stauning alper (72°n), and made first ascents of merchiston tinde, dunot tar bjerg and tantallon spids. a crossing was made of the south stauning alper from alpefjord to sydkap, and the first traverse of the central stauning alper from gully gletscher to bersærkerbræ was completed via col major (majorpasset). limited glaciological 63 work was carried out on lower sefström gletscher. a few climbs were also made west of alpefjord. many glaciers and mountains were named, and most names have approved status. slesser’s names were mostly given after scottish castles (bennet 1959; slesser 1959, 1964a, b). 1958 carlsberg foundation scoresby sund expedition botanical and biological studies were carried out by two parties, supported financially by the carlsberg foundation. co-operation with lauge koch’s expedition provided air transport facilities. parties visited many localities between gåseland (70°n) in the south, and geographical society ø and ella ø (73°n) in the north. 1958–59 grønlands geologiske undersøgelse (ggu) expeditions to kap stosch in both 1958 and 1959, small parties, led by svend e. bendix-almgren and with support from grønlands geologiske undersøgelse (ggu), visited the kap stosch region (74°03´n). geological and palaeontological collections were made. 1959 ‘operation groundhog’: j.m. hartshorn an investigation of ice-free sites for emergency air craft landings was carried out between 70° and 74°n in east greenland by scientists of the united states air force cambridge research center and the united states geological survey under us air force con tract. the six-man scientific party, led by j.m. hartshorn, was based on the icebreaker uss atka and operated in the region from 15 august to 10 sep tember. of numerous potential sites selected from studies of aerial photographs, many were inspected briefly by helicopter, and a few were mapped and marked out. special attention was given to sites around scoresbysund, southern ymer ø and storelv (hartshorn et al. 1961). helge larsen accompanied the party and made archaeological observations. no new place names are recorded in their official report. 1959–60 tristan jones voyage with the yacht cresswell tristan jones made a single-handed sailing voyage to east greenland via iceland in his converted wooden lifeboat cresswell. after reaching scoresby sund (70°n) in august 1959, he sailed northwards along the coast almost to kap bismarck (76°42´n), but was trapped in the pack ice, and drifted with the ice down to the latitude of scoresby sund, where he met the gustav holm. refusing an offer of a lift to iceland, tristan jones with the cresswell overwintered at sydkap from october 1959 to may 1960. after leaving east greenland, jones sailed for spitsbergen, where he was again caught in pack ice and the cress well was lost (jones 1979, 1983). 1959–61 american glaciological expeditions fred pessl and norman p. lasca carried out glaciological studies around the head of mesters vig (72°n) in 1959 and 1961, funded from american sources, and with local help from nordisk mineselskab (pessl 1962). 1959–64 geodætisk institut (geodetic institut) aerial photography aerial photography was carried out for the danish geodetic institute over large areas of north and east greenland from a base at station nord (81°36´n) in northern kronprins christian land. vertical photo graphs at a scale of 1:50 000 were obtained for the en tire region north of 76°n, while a number of oblique routes were flown in the scoresby sund region in 1961. 1960 ‘operation groundhog’: s.m. needleman the united states air force cambridge research cen ter and the united states geological survey under air force contract carried out scientific studies and investigations of emergency aircraft landing strips at centrumsø (80°10´n), culminating in test landings by a canadian air force c-119 and a us air force c130 hercules aircraft. the scientific parties, led by s.m. needleman, received some support from the us army ‘operation lead dog’ working on the ice cap nearby (needleman 1962). two new names (græselv and grottedal) later came into use in the area covered by this volume. three of the scientists, led by w.e. davies, worked for part of the summer in peary land, northernmost greenland. 64 1960 british east greenland expedition: john hunt john hunt led a party of 38, including 21 boys, on a largely climbing expedition to the stauning alper (c. 72°n). several first ascents were made around ber særkerbræ, including that of the hjørnespids (slesser 1961, 1964a, b). from alpefjord, reached with the motor boat, polypen, the party traversed via spærre gletscher and duart gletscher into the southern stau ning alper, where several mountains were climbed around bjørnbo gletscher. about 12 mountains and 14 glaciers were named, and nearly all have approved status. following the system introduced by malcolm slesser in 1958 the mountains were named after scottish castles, while glaciers were named after planets or constellations of stars (jackson et al. 1961; hunt & sugden 1962; slesser 1964a, b). some glaciological and ornithological observations were also made. 1960 usaf aerial photography a small number of vertical aerial photography routes were flown by the united states air force (usaf) over parts of the scoresby sund region (70°–72°n). modern scientific investigations, adven ture and sporting expeditions 1961–2008 1961 bangor junior mountaineering club expedition: m.k. lyon a nine-man expedition led by m.k. lyon explored the region around the schuchert gletscher and stor gletscher in the southern stauning alper (71°55´n), making several first ascents, the most notable being royal peak. one man had a bad accident and was flown out to iceland. a brief account of the expedition is given in bennet (1972). the names given for the peaks climbed show no clear system, and none have acquired approved status. 1961 junior mountaineering club of scotland expedition: james clarkson encouraged by the reports of john hunt’s 1960 expedition, a nine-man group led by james clarkson explored the bjørnbo gletscher system in the southern stauning alper (71°40´n) making 24 first ascents. a return crossing of the range from bjørnbo glet scher to alpefjord via spærregletscher was also made (clarkson 1962, 1964). following earlier usage many of the peaks were named after scottish castles, while other peaks and 12 glaciers were named after heavenly bodies. proposals to authorise the names were made, but in contrast to its earlier practice the place name committee now declined to accept large num bers of ‘foreign-sounding’ names within the stauning alper. [place name committee archive.] 1961 cambridge east greenland expedition: russel marris a party of six led by russel marris visited the fleming fjord and mestersvig areas (71°30´–72°20´n), ma k ing ornithological and biological observations (mar ris & ogilvie 1962; hall 1964). a total of 569 barnacle geese and six pink-footed geese were ringed. 1961–1962 leicester university east greenland expeditions: geoffrey halliday these expeditions carried out a varied programme of botanical, geological and zoological work, and flew into mestersvig from iceland by chartered aircraft. 1961 – geoffrey halliday led a party of 12, mainly from leicester university (england), to the region west of mestersvig and between 8 july and 9 sep tember visited forsblad fjord, alpefjord and furesø (72°n). from the head of furesø, reached by boat, a traverse was made via jomfrudal to the coast of nordvestfjord. botanical, zoological and geological observations were made. two members also reached the highest point of the ice cap north of furesø, approached via schaffhauserdal (halliday 1962, 1963). five unapproved names are recorded. 1962 – geoffrey halliday continued his botanical studies in east greenland from 18 july until 10 sep tember, leading a five-man group to the kong oscar fjord region (72°10´n). investigations were concentrated in the area around mestersvig, the coast of the northern stauning alper and southern traill ø (halliday 1963). 1961–84 arktisk minekompagni (arctic mining company) arktisk minekompagni was a consortium with 50% interests held by respectively nordisk mineselskab 65 and amax (american metal climax inc.), formed to undertake investigations of the molybdenum deposit at malmbjerg (72°n). an exclusive consession to mine and ship molybdenum was granted in 1961, but the concession was suspended in 1984 in association with negotiations over oil exploration rights in jame son land (see below). extensive drilling of the prospect was carried out in 1961 and 1962, the 67 drill holes bringing the total length drilled up to 20 km. reserves of close to 200 million tons of ore with 0.25% molybdenum sulphide were proven (thomassen 2005b). a small mining ‘town’ of wooden barracks was built for the drilling crews on the moraines just south of the deposit, but the site was cleared in the 1980s. the situation of the deposit, surrounded by glaciers, and its relatively low grade has so far hindered exploitation, but following dramatic price rises for molyb denum, new investigations were initiated in 2005. 1962 oxford university expedition to east greenland an eight-man scientific party led by d.e. sugden and b.s. john visited pingodal and schuchert dal in jame son land, and oxford gletscher in the southern stau ning alper (c. 71°30´n). they undertook geo morp h ological, ornithological and botanical studies. the party flew in via mestersvig, and sailed home from scoresbysund with the kista dan (john & sugden 1963; worm 1963). four new approved names resulted from the expedition’s work, including the name oxford gletscher, while several unapproved names have appeared in ornithological reports (hall 1963, 1966). 1963 geodætisk institut (geodetic institute) expedition to scoresby sund surveying teams from the danish geodætisk institut carried out triangulation in the inner scoresby sund region in 1963, supported by the ships tycho brahe and ole rømer. about 10 new names were proposed for various features, all of which are approved. 1963 trinity college east greenland expedition: k.c. campbell a party of 10 led by k.c. campbell, mainly from trinity college, dublin, carried out botanical and ornithological studies in hurry inlet, carlsberg fjord and the jameson land coast of hall bredning (70°30´–71°30´n). some of their equipment was airdropped onto jameson land from their dc-4 aircraft (campbell 1964). elio pampanini flew a beechcraft bonanza aircraft to the region in august to assist in the evacuation of the expedition. 1963 cambridge university east greenland expedition: colin f. knox this 12-member climbing expedition from cam bridge university (england) was led by colin f. knox, and concentrated its activities in the region of gullygletscher and sefström gletscher in the stau ning alper (72°n). they were assisted by an airdrop of food and equipment onto sefström gletscher at the beginning of the season. a total of 25 first ascents were claimed, including c.f. knox tinde, sne top pen, pembroke kuppel, korsspids and canta brigia tinde, all over 2700 m high (roschnik 1964; knox 1964a, b). several long traverses were also made, and the expedition is generally considered to have been one of the most successful to have visited the stauning alper (bennet 1972). most of their named peaks commemorate cambridge colleges, or have associations with cambridge, and were subsequently ap proved in danicised form. one of their peaks, grandes jorasses, was subsequently renamed c.f. knox tinde following the death of knox in the french alps in 1964. some glaciological work was carried out on the lower sefström gletscher. 1963 la spedizione italiana, g.m.’63 (italian expedition to the stauning alper): guido monzino the italian climber guido monzino led a group of 14 italians to the bersærkerbræ region of the stauning alper (72°n); the group could not reach their original goal around petermann bjerg due to the presence of winter ice in the fjords. five camps were set up on bersærkerbræ, and the second ascent of glamis borg (cima di granito) was made by a new route (fantin 1969; bennet 1972). 1963 british east greenland expedition: russel marris russel marris led an eight-man party which visited the ørsted dal area (71°47´n) to make ornithological 66 observations (hall & waddingham 1966). one of their main objectives was to ring barnacle geese. some geological observations were made around pingel dal. 1963 imperial college east greenland expedition: m.h. key a climbing group from imperial college (london, england) led by m.h. key visited the stauning alper (72°n), concentrating on the peaks around the ber særkerbræ. some glaciological and geological studies were also made. of the 24 mountains climbed, 15 were first ascents. the names proposed for their peaks were all given after london boroughs, because all the members of the group were from a london college. however, while some attempt was made to seek official approval of their names, the formalities were never concluded. accounts of the expedition are given by key (1964) and watson (1964). 1964 la spedizioni italiana, g.m.’64 (italian expedition to the stauning alper): guido monzino guido monzino returned to the stauning alper (72°n) with a party of 20 climbers; from mestersvig they travelled to alpefjord by inflatable boat. two peaks on the south side of vikingebræ were climbed, cima est and cima oest, and the second ascent of dansketinde was made by a new route (bennet 1972). a 1550 m peak south of kap peterséns was also climbed. 1964 expedition des academischen alpen clubs zürich in die stauningsalpen (academic alpine club zurich expedition to the stauning alper): a. hofmann a party of 10 swiss climbers led by a. hoffman made five first ascents in the syltoppene area of the northernmost stauning alper, subsequently moving by inflatable boat to the sefström gletscher region (72°n) where a further four first ascents were made. finally they moved to the spærregletscher area and climbed at least another eight peaks (meinherz 1965). some reports record a total of 21 first ascents (fantin 1969). none of their place names have acquired official status. 1964 daneborg ornithological expedition a danish three-man ornithological expedition made observations in the daneborg region (74°18´n) be tween mid-april and mid-july. from a base at dane borg weather station, where a landing was made on the sea-ice with a dc-3 on 18 april, journeys were made northwards as far as germaniahavn and linde man fjord, and westwards to revet (christensen 1965, 1967; rosenberg et al. 1970). 1965 oxford university expedition to east greenland: j.c. rucklidge a geological expedition of six men from oxford uni versity (england) led by j.c. rucklidge sailed to scoresbysund with the thala dan, and then crossed scoresby sund with the settlement boat entalik to reach their working area near kap brewster, the ba salt region on the south side of scoresby sund (70°n). an advance base was established near the front of torvgletscher, from which journeys were made to the upper reaches of the glacier, and pind svinet was climbed (rucklidge 1966). the expedition was pick ed up by the entalik on 4 september, but because of bad ice conditions was forced to abandon much of their equipment (rucklidge & brooks 1966). no new place names are recorded. [rgs report archive.] 1966 cambridge expedition to east greenland: russel marris russel marris and a.m.f. webbe made botanical and ornithological observations in the region between mestersvig and daneborg (72°–74°30´n), with especial reference to the barnacle geese (marris & webbe 1970). 1966 deutsche grönland-expedition in die staunings-alpen (german expedition to the stauning alper): karl m. herligkoffer a six-man party led by karl m. herligkoffer had originally intended visiting the peary land region of north greenland, but was frustrated by lack of aircraft fuel at mestersvig preventing them from con tinuing their journey. instead they combined forces with a four-man group from munich which had sailed to mestersvig with the nella dan, and turned their attention to the nearby stauning alper (72°n). about 30 first ascents were claimed in the region around the 67 heads of spærregletscher, roslin gletscher and borg bjerg gletscher (herligkoffer 1967); some were probably second ascents (bennet 1972). an attempt was made to gain approval of the names for their peaks, which were mainly given for german towns or localities, but their localities were said at the time not to be sufficiently precise. most peaks have since been located on modern maps (bennet 1972); see also map 5. 1967 grønlands geologiske undersøgelse (ggu) international expedition to kap stosch a nine-person group of danish, swiss and american geologists, with support from grønlands geologiske undersøgelse (geological survey of greenland, ggu), visited the kap stosch region (74°03´n) to undertake geological and palaeontological studies of permian and triassic rocks. 1967 ohio state university expedition a two-man american party, john gunner and dave par rish, made a visit to the inner fjord region of scores by sund (70°n), including a two-week walking trip from inner føhnfjord along hjørnedal to the interior of gåseland around gnejssø. 1967 lambert land search expedition j.l. christiansen and n. preben-andersen visited lambert land (79°15´n) by catalina in mid-august to search for traces of mylius-erichsen and høeghagen, two of the three members of the 1906–08 dan mark-ekspeditionen who had died in 1907. nothing significant was found. 1967 berchtesgaden expedition to the stauning alper four german climbers from berchtesgaden (ger many), visited the area west of spærregletscher (72°n), making 13 first ascents. their highest peak was schnee kuppel, 2640 m high. their names were apparently given for german localities and notable mountaineers, but none have approved status. sum mary accounts of the expedition are to be found in bennet (1972), fantin (1969) and hoff (1979). 1967 spedizione sci-alpinistica italiana in groenlandia (italian expedition to the stauning alper): toni gobbi this italian climbing party of 12 led by toni gobbi was most interested in ski-mountaineering, and visited the stauning alper (72°n) from mid-june. a num ber of climbs were made from the bersærkerbræ including dunottar bjerg and kensington, and a first ascent was made of panoramic peak (fantin 1969; bennet 1972; hoff 1979). 1967–69 geodætisk institut (geodetic institute) surveying and aerial photography triangulation was carried out in 1967 on the blosse ville kyst (69°n), mainly south of d’aunay bugt, based on the motor cutter ole rømer. in 1968 and 1969 a variety of surveying objectives were carried out from the boats ole rømer and tycho brahe, to increase the detailed triangulation network and density of fixed points. the 1968 work included a survey of schuchert dal and malmbjerg. aerial photography was carried out in 1968 and 1969 of parts of the scoresby sund region (70°–72°n), although it was considerably hindered in 1969 by poor weather. 1967–72 grønlands geologiske undersøgelse (ggu) scoresby sund expeditions a series of major expeditions by the geological sur vey of greenland (grønlands geologiske under søgelse: ggu) to the scoresby sund region were led by niels henriksen. they had as their principal ob jective the systematic geological mapping of the region 70°–72°n, to be published as 1:100 000 and 1:500 000 scale map sheets (henriksen 1986). a twoman reconnaissance expedition in 1967 based on a small cutter jytte visited the entire fjord system, and provided a logistical and geological background for planning of the subsequent major expeditions. 1968 – the 31-member expedition sailed to scores by sund aboard the martin karlsen (formerly the kista dan), which functioned as a floating base throughout the summer for two helicopters serving 12 geological teams. activities were concentrated in the inner parts of nordvestfjord, and in northern jameson land. 1969 – the magga dan was the expedition ship, and carried the party of 38, mainly scientists, to 68 scores by sund, as well as acting as base ship for the two helicopters (fig. 22). the 15 geo logical teams wor ked in the southern stauning alper, renland, jame son land and southern liver pool land. 1970 – the expedition numbered 43, including 16 geological teams, and operated with two helicopters from the base ship perla dan. the main working areas were in renland, milne land and areas west of rødefjord, with five groups working on the mesozoic rocks of eastern milne land and jameson land. 1971 – this 44-member expedition operated from a tent base camp at the head of hurry inlet, the 16 teams of geologists being served by three helicopters. the working areas were liverpool land and jameson land, with two teams working on the basalts south of scoresby sund. the norwegian sealer brandal transported fuel and supplies to the region, and also supported a geophysical group working in scoresby sund. a three-man ggu group carried out an aeroradiometric survey of selected areas of east milne land and schuchert dal using a dornier 28 aircraft; these studies continued in subsequent years (see ‘1971–77 ggu/aek expeditions to east greenland’ below). 1972 – the last year of the scoresby sund expeditions worked out of a land base at hjørnedal in føn fjord, the 44 participants being served by three helicopters and a pilatus porter stol (short takeoff and landing) aircraft. working areas for the 14 teams were on southern milne land, gåseland, and the basalt areas along the south side of scoresby sund. the detailed mapping and exploration in areas only scantily investigated by earlier expeditions led to approval of 70 new place names for large and small features, their derivations being as diverse in character as the numerous geologists who proposed them. 1968 scottish expedition to the stauning alper this seven-man expedition traversed from mesters vig (72°13´n) overland to the central stauning alper, making the first crossings of passes between edinbræ and schuchert gletscher, and between storgletscher and grantagletscher. two first ascents were made of peaks on the north side of sefström gletscher, as well as the third ascent of sefström tinde (bennet 1969, 1972). 1968 graham tiso’s east greenland expedition graham tiso led a five-man climbing party to the gully gletscher region of the stauning alper (72°n). the third ascent of norsketinde was made by a new route, after which the party crossed alpefjord to climb in eastern nathorst land around trekant gletscher (hill 1969; bennet 1972). 1968 nordost grönland expedition (german north-east greenland expedition): hermann huber hermann huber led a four-man german climbing ex pedition to the vikingebræ region of the stauning alper (72°n). several first ascents were made, including dreispitze and högspids (fantin 1969; bennet 1972). 1968 university of dundee scoresby land expedition: ian h.m. smart iain h.m. smart led an eight-man group with climb69 fig. 22. the magga dan was a polar expedition ship built for the j. lauritsen shipping company. in 1969 it was the expedition ship for ggu’s summer expedition to the scoresby sund region. fitted with helicopter platforms at the front and rear it functioned as a floating base. ggu’s cutter jytte seen moored to the ship to the right provided local transport for geological parties. ing and scientific objectives to the southern stauning alper and pingo dal (71°40´). from mestersvig the par ty walked via mellempas to malmbjerg, and thence to pingo dal where pingos were studied and surveyed until the end of july; pingos are ice-cored conical mounds found on braided river plains (see also fig. 70). the climbing group found a new route to the peaks at the head of roslin gletscher, and made nine first ascents (bennet 1972). smart carried out studies of arctic terns on the menander øer. none of the names given for their peaks have been approved. one of these, dreverspids, commemorates the patrons of the expedition, james and harald drever. an associated party of six from edinburgh university and the university of dundee, including george and irene waterston, carried out ornithological and biological studies between antarctic havn and mestersvig (waterston & waterston 1969). [rgs report archive.] 1968 expédition française au groenland nord-est (french expedition to north-east greenland): claude rey a large climbing expedition of 16 men and women led by claude rey sailed to the head of dammen by rubber dinghy, and explored the area around prinses se gletscher, at the west margin of the stauning alper (72°n). nineteen first ascents were made of the high peaks on both sides of the glacier, including several long climbs and traverses (georges & rey 1969; ben net 1972). 1968 womens’ east greenland mountain eering expedition: joan busby a five-member women’s expedition led by joan busby made a number of climbs in the bersærkerbræ region of the northern stauning alper (hoff 1979). 1968 ornithological studies: russel & david marris the brothers russel and david marris visited the scoresby sund region (70°–72°n) in 1968, travelling in the fjords by small boat. their activities led to the approval of four names, mainly with botanical ori gins. [place name committee archive.] 1968–70 cambridge greenland expeditions: peter f. friend a series of geological expeditions led by peter f. friend of cambridge university (england) visited the region 71°30´–74°30´n, with the main purpose of in vestigating the devonian sandstones. each year the party arrived at mestersvig by chartered aircraft, con tinued to ella ø by catalina aircraft or boat, and sub se quently used inflatable boats for transport through out the fjord system. occasional use was made of chartered helicopters to reach inland areas. the parties num bered 11 in 1968, 10 in 1969 and 12 in 1970. three new place names were introduced in the course of their studies (friend et al. 1983). 1968–75 east greenland expeditions: keith j. miller keith john miller [1932–2006] was a mechanical en gineer of world standing, and an enthusiastic mount aineer. his expeditions to east greenland often com bined scientific activities with climbing, and were partly used to develop radio echo-sounding techniques for measuring the thickness of ice in glaciers. 1968 – keith miller led an eight-man party from queen mary college, london (england) to the bersærkerbræ region of the stauning alper (72°n). climbing groups made the first ascent of bersærker tinde, the second ascent of hjørnespids and the third ascent of dansketinde. glaciological work was car ried out on besærkerbræ. miller fell into a crevasse and was evacuated to reykjavik for treatment, while another member of the party (tom hird) fell into a melt-water stream on the glacier and was lucky to escape with minor injuries (bennet 1972). 1970 – this ten-man chiefly scientific expedition led by keith j. miller of cambridge university (uk) flew by helicopter from mestersvig to roslin glet scher (71°48´n), where the british royal air force had parachuted in supplies and equipment. glacio logical studies included echo-sounding experiments to determine the thickness of the glacier ice. three peaks over 2000 m high were climbed, two of them first ascents. 1972–73 – a 12-man party led by keith j. miller from cambridge university continued their studies on the roslin gletscher (71°48´n). their main pro jects included testing a thermal ice probe, and radio echo-sounding of ice thickness. in 1972 they co-oper70 ated with a two-man imperial college greenland ex pedition and with the cambridge schuchert ex pedition. six peaks were climbed at the end of the summer, including two first ascents. in 1973 a fourman party continued the work. 1975 – keith j. miller led a four-man group from cambridge university (england) to the stauning alper, that made a spectacular and very long (250 km) n–s traverse of the range from kap peterséns in the north to sydkap in the south, including crossing two new passes. the return to mestersvig was made via schuchert dal (miller l976). [rgs report archive.] 1969 spedizione sci-alpinistica italiana in groenlandica (italian ski-mountaineering expedition to the stauning alper): toni gobbi a 13-member italian climbing party led by toni gobbi visited the bersærkerbræ region of the stauning alper (72°n), making several first ascents (bennet 1972). their prime interest was ski-mountaineering; a particularly fine ski traverse was made by one group via skelgletscher, schuchert gletscher and sefström gletscher to alpefjord, returning via gullygletscher and majorpasset (col major). 1969 zoogeographical investigations: christian vibe & ivar silis christian vibe and ivar silis carried out studies of polar bear and musk ox in the daneborg and clave ring ø areas (74°20´n). 1969 norwegian musk-ox expedition: john j. teal john j. teal of the university of alaska was leader of an expedition aboard the harmoni which visited kejser franz joseph fjord in search of musk oxen. twenty-five young musk oxen were captured, and taken back to norway for release in the bardu dis trict. 1969 watkins bjerge expedition: a.j. allen a.j. allen led a six-man anglo-danish expedition to the scoresby sund region whose aim was to reach the watkins bjerge (69°n) from the north. the party flew into scoresbysund in early july, but the break-up of the fjord ice frustrated their planned sledge journey, and they eventually reached danmark ø by boat. on 22 july they were lifted by helicopter to sydbræ. a long journey across geikie plateau brought them to within 30 km of their goal, but very poor weather led to a retreat to innermost gåsefjord, where the party was picked up by the entalik on 28 august. [rgs report archive.] 1969 international mount mikkelsen expedition: malcolm slesser the objective of this four-man expedition led by c.m.g. (malcolm) slesser was to climb ejnar mik kelsen fjeld (69°n), 40 km inland from the blosseville kyst. three of the party flew to scoresbysund, while the fourth (carlos ziebell) reached gurreholm by air from mestersvig then walked the rest of the way to scoresbysund. the party sailed from scoresbysund in an open boat southwards to kap brewster, and down the blos se ville kyst as far as the south-east point of turner ø. however, because of delays due to storms and difficult ice conditions they succeeded only in climbing a few minor peaks near the coast (smart 1970; slesser 1970). hot springs in rømer fjord were investigated. six place names, mainly with scottish associations, were approved. [rgs report archive.] 1969–71 hans meltofte ornithological observations, danmarkshavn while employed at danmarkshavn weather station (76°42´n) from april 1969 to april 1971, hans meltofte made regular ornithological observations (meltofte 1975). observations were concentrated in the vicinity of the station, but sledge journeys were also made northwards to kap amélie, and westwards to annekssøen, sælsøen, ålborghus and rechnitzer land. more than 500 birds, mostly snow buntings, were ringed. seven names reported by meltofte as in use by personnel at the station were subsequently formally approved. 1970 british expedition to ejnar mikkelsen fjeld: andrew ross andrew ross led a party of four which made the first successful ascent of ejnar mikkelsen fjeld (68°53´n) in the watkins bjerge. the approach was made from scoresbysund down the blosseville kyst in a large open boat. on the return voyage along the coast, the 71 party was caught in bad weather, lost their fuel sup plies, and were rescued by the perla dan (the ggu expedition ship) at søkongens bugt (68°40´n; ross 1971). 1970 scottish expedition to the stauning alper david bennet and malcolm slesser climbed together in the stauning alper (72°n), making a new and easier route on the bersærkertinde, and the first ascent of a small rock peak to its east (bennet 1972). 1970 st. andrews university east greenland expedition: r.m. nisbet this climbing expedition from st. andrews uni ver sity (scotland) was led by r.m. nisbet, and climbed seven peaks in north-east nathorst land around schaff hauserdalen (72°20´n). nisbet broke a leg in an accident, and was evacuated by helicopter (bennet 1972). 1970 münchner grönland-fahrt (german climbing expedition to nathorst land): wolfgang weinzierl this german climbing expedition led by wolfgang weinzierl visited north-east nathorst land (72°n), and made seven first ascents around trekantgletscher and one in the stauning alper (weinzierl 1971). the brief report is confusing as directions are misleading (e.g. trekantgletscher is said to be ‘east’ of alpefjord whereas it is to the west). the peaks are also very difficult to locate as the report has no map. 1970 ladies’ scottish east greenland expedition: helen steven a party of 12 ladies led by helen steven climbed in the stauning alper and nathorst land (72°n). five ascents were made west of bersærkerbræ, including a repeat of the bennet/slesser route on the bersærker tinde, and four climbs (three first ascents) in nat horst land (bennet 1972; hoff 1979). 1970 expédition française au groenland nord-est (french expedition to the stauning alper): claude rey a climbing group led by claude rey sailed from mesters vig to alpefjord and made five climbs in the vikingebræ region (72°10´n). these included the fourth ascent of norsketinde, and the first ascent of mythotinde (bennet 1972). 1970 university of dundee scoresby land expedition this 14-man university of dundee (scotland) expedition to the central and southern stauning alper (72°n) was organised as four groups, mainly operating independently. three groups subsequently com bined to carry out glacier exploration and mountaineering in the southern stauning alper. seven or eight peaks were climbed, mainly first ascents, two of which received unofficial names – taurobjerg and boulderbjerg (bennet 1972). hydrological and biological studies were also made. a boat journey was made by one party to the bjørneøer and into nordvestfjord as far as nordbugt. [rgs report archive.] 1970–73 swedish expeditions to east greenland these swedish expeditions were active between fle ming fjord and hold with hope (71°40´–74°n), and were primarily concerned with quaternary geology and ornithology (hjort 1976). 1970 – christian hjort and three others visited the kong oscar fjord region. 1971 – a party including christian hjort visited the area around mestersvig, lyell land, ella ø and the east coast of geographical society ø. 1972 – visits were made to fleming fjord, traill ø and kempe fjord. 1973 – c. hjort and j. mikaelsson visited the hold with hope and hudson land region. 1971 radley college east greenland expedition: g. treglown a party of six from radley college (uk) led by g. treglown flew into mestersvig by british royal air force hercules at the end of july. ornithological studies were made from camps near mestersvig, on traill ø and ella ø (72°–73°n; hardy 1979). [rgs report archive.] 72 1971 grumman ecosystems aerial photography vertical aerial photography was carried out by grum man ecosystems corporation for greenarctic con sortium, over selected areas between 74°n and 76°n in east greenland. greenarctic consortium was a large prospecting company with interests in the da nish and canadian arctic. 1971 university of lancaster expedition to the southern stauning alper: harry pinkerton a three-man university of lancaster (england) expedition led by harry pinkerton to the southern stau ning alper (71°40´n), was later joined by two mem bers of the ‘1971 northern universities east green land expedition’. four ascents were made around bjørnbo gletscher, three of them first ascents (bennet 1972; pinkerton 1972). 1971 expédition française au groenland nord-est (french climbing expedition to north-east greenland): claude rey a small french climbing group led by claude rey visited the vikingebræ region of the stauning alper (72°10´n). among other climbs, the first ascent was made of a peak north of helvedespas (bennet 1972). 1971 american east greenland expedition: george wallerstein george wallerstein led a party of six american clim bers that intended to make an attempt on ejnar mikkelsen fjeld (68°53´n) from the north. the party failed to reach their goal, but made a reconnaissance of sydbræ (70°n), and climbed three minor peaks in milne land (liska 1972; hoff 1979). they had great problems returning to scoresbysund when their boat was trapped by pack ice on the shore of jameson land. 1971 british expedition to the roscoe bjerge, liverpool land: malcolm slesser c.m.g. (malcolm) slesser led a party of six to southern liverpool land, carrying out ski-mountaineering and making nine first ascents in the roscoe bjerge, liverpool land (70°39´n; slesser 1972). 1971 northern universities east greenland expedition: geoffrey halliday a british, largely scientific, party of up to nine members led by geoffrey halliday visited the scores by sund region to carry out botanical, ornithological and geological studies. supplies were air-dropped at scoresbysund, gurreholm and nordbugt. one party flew into scoresbysund and worked in southern liver pool land (70°40´n). a second party flew into mestersvig and walked to gurreholm, from where a boat journey was made along nordvestfjord to nord bugten and flyverfjord (71°33´n). several long walks were made inland from nordbugt, in hinks land and along edvard bay dal. two members joined a climbing group from the ‘1971 university of lan caster expedition’ that ascended several peaks in the bjørn bo gletscher region, including three first ascents (bennet 1972; pinkerton 1972). five names in the inner reaches of nordvestfjord, given as botanical reference localities, were subsequently approved. some are given for plants, two others leeds and lan caster universities. [rgs report archive.] 1971–72 atlantic richfield oil company (arco) arco in association with nordisk mineselskab carried out geological studies over an extensive region in east greenland. up to three helicopters were used to transport geological teams and equipment, and these gave occasional assistance to the various sports expeditions in the region. 1971–77 ggu/aek expeditions to east greenland co-operation between grønlands geologiske under søgelse (geological survey of greenland: ggu) and the atomenergikommissionen (danish atomic ener gy commission: aek) led to an extended series of activities, including aero-radiometric surveys, uranium prospecting and ground-based studies of radioactive anomalies. 1971 – an aerial gamma spectrometric survey was carried out in july and august between scoresby sund (70°n) and hold with hope (74°n) using a dornier 28 twin-engine aircraft. follow-up ground investigations were initiated. 1972 – follow-up ground investigations of anomalies were continued (nielsen & løvborg 1976). 1973 – a 15-person group was based at stordal. 73 this undertook a detailed airborne geophysical pro gramme in the same region as in 1971, and extended coverage to 76°n using a britten-norman islander aircraft. this included radiometric and aeromagnetic surveys with an average ground clearance of 100 m (nielsen & larsen 1974). 1974 – a further, large group based at stordal continued systematic geophysical prospecting using a britten-norman islander. ground prospecting was carried out of anomalous localities, assisted by a helicopter carrying a scintillometer. a geochemical samp l ing programme of stream waters and sediments was commenced. 1975 – the uranium prospecting programme based at stordal was continued, with detailed helicopter-supported geological and radiometric investigation of radioactive anomalies, and geochemical sampling. 1976 – a large group continued detailed investigations of radioactive anomalies, and continued geo che mical sampling, assisted by a helicopter. 1977 – the last of the stordal-based expeditions completed the geochemical water and sediment sampling programme, and follow up studies of the gamma-spectrometer work. detailed work was car ried out on previously detected anomalies (steenfelt & nielsen 1978). from 1974 onwards, small groups of ggu geologists took advantage of the stordal facilities to carry out general mapping projects – see ‘1974–79 grøn lands geologiske undersøelse (ggu) mapping pro jects in east greenland’ below. 1972 university of dundee north-east greenland expedition: r.m.g. o’brien r.m.g. o’brien was leader of an 11-member uni versity of dundee (uk) expedition which carried out ornithological and zoological observations in andrée land, ymer ø and the mestersvig area (72°–74°n; sum mers & green 1974). the expedition was assisted by british royal air force (raf) air drops made at kap peterséns and renbugten. travel was by inflatable boat and on foot, and a long traverse was made from ren bugten via djævlekløften to grejsdalen in an drée land. in addition, three peaks were climbed at the head of haredalen on the west side of isfjord. [rgs report archive.] 1972 h.w. tilman’s voyage with the seabreeze h.w. (bill) tilman took to sailing in 1955 as a means of reaching unclimbed mountains, and made three voyages to west greenland and two to the ammas salik region of east greenland in his pilot cutter mischief. he later made three attempts to reach scoresby sund with the seabreeze, a 49-foot bristol channel pilot cutter, the most successful in 1972 when he came to within a few kilometres of kap tobin (71°24´n). his earlier voyages in 1969 and 1971 were, as in 1972, frustrated by pack ice at the mouth of scoresby sund. the seabreeze ran aground and foundered south of ammassalik in 1972 on her way home (tilman 1974). 1972 knud lauritzen, summer cruise during august of 1972, knud lauritzen, owner of the danish j. lauritzen shipping company, sailed through parts of the scoresby sund fjord system (70°–72°n) in his motor yacht bamsa dan. this included a circuit of milne land and a visit to the ggu base camp at hjørnedal. 1972 cambridge schuchert expedition: f. alayn street six ladies from the geography department of the university of cambridge (england) led by f. alayne street undertook botanical and glaciological studies near the terminal moraines of roslin gletscher in schuchert dal (71°48´n). they were assisted by bri tish royal air force (raf) air-drops, and an occasional helicopter lift from the atlantic richfield company (arco). [rgs report archive.] 1972–73 geodætisk institut (geodetic institute) surveying and aerial photography the danish geodetic institute (geodætisk institut) continued in 1972 their improvement of point control on the blosseville kyst with a party based on the ole rømer. an attempt to fix the position of the gronau nunatakker south-west of gåsefjord was unsuccessful. vertical aerial photography was flown over large parts of the scoresby sund region in 1972. in 1973 aerial photography coverage was extended northwards to 74°30´n, but was brought to an untimely end by the crash of the aircraft at mestersvig airfield with the death of the pilot and one of the surveyors. 74 1972–73 imperial college greenland expedition: peter w. chaplin 1972 – peter w. chaplin and richard a. carter from imperial college (london, uk) accompanied the cambridge stauning alper expedition to the ros lin gletscher region (71°48´n) of the southern stau ning alper, where they maintained meteorological records and made a plane table survey of the routes of echo-sounding traverses. [rgs report archive.] 1973 – a party of four led by peter w. chaplin revisited the roslin gletscher (71°48´n), where the stake lines of the 1972 expedition were re-surveyed, despite difficulties with heavy snow. new lines were surveyed at the front of storgletscher. the party walked south as far as sydkap, before returning to mestersvig on foot (chaplin et al. 1976). [rgs report archive.] 1973 sheffield university geological expedition to mestersvig: charles downie charles downie of sheffield university (uk) led a four-man geological group to the mestersvig area (72°n), whose objectives included sampling the meso zoic sequence at antarctic havn and on traill ø with particular reference to the oil resource potential. 1973–75 swedish scoresbysund expedition: magnus elander magnus elander made two summer expeditions to east greenland to undertake environmental studies of trace amounts of poisons in birds and animals. in 1972 he visited mestersvig and scoresbysund. in 1975 work was carried out from bases on rathbone ø (70°40´n) and in hurry inlet. [dpc report archive.] 1973 hurry inlet expedition: r.m. sykes & s.r.a. kelly r.m. sykes and s.r.a. kelly visited the hurry inlet region (70°40´n), making stratigraphical observations and palaeontological collections (sykes & callo mon 1979). 1973–75 nederlandse groenland expeditie (dutch greenland expeditions) two or three-man ornithological expeditions from dutch universities and research organisations visited jameson land and southern liverpool land three years in succession. their main study was the ecology of the long-tailed skua, and additional studies were made of waders that winter in or migrate through holland. in 1973 their base camps were on rathbone ø and at kap stewart, in 1974 at kap stewart and near kærelv, and in 1975 at kærelv (korte et al. 1981). 1973–75 de danske isbjørneekspeditioner (the danish polar bear expeditions): christian vibe christian vibe led a series of expeditions to east greenland to study and mark polar bears in their main breeding area, the fjord region between 69° and 78°n latitude, most of which lies within the borders of nordøstgrønlands nationalpark established in 1974. the expeditions were supported by the danish natural science research council (snf) and the ministry for greenland. activities took place mainly in the spring, in 1973 using snowscooters and a small cessna 185 aircraft, in 1974 helicopter and small aircraft, and in 1975 when activities extended into the pack ice off the coast, the norwegian sealer polar star and helicopter. scoresbysund, mestersvig air field, daneborg and danmarkshavn were used as sup port bases. the observations suggested there was a resident population of about 200 bears in the region, with regular additions to the population drifting in with the pack ice from spitsbergen (vibe 1982). 1974 nordøstgrønlands nationalpark (north-east greenland national park) eske bruun (1966) had argued strongly for the establishment of a national park in northern east green land at a time when the ‘østgrønlands traktat’ was about to expire in 1967. christian vibe was an enthusiastic supporter of the idea as a result of his wideranging studies of musk oxen and polar bears (vibe 1967, 1971, 1982, 1984). in 1974 these ideas came to fruition when greenland’s first national park was established; after expansion across north greenland in 1988 ‘nordøstgrønlands nationalpark’ became the largest national park in the world; it has sometimes been referred to as ‘nordog nordøst grønlands na tional park’ (north and north-east green land na tional park). the park incorporates the land areas of northern east greenland with a southern boundary at approximately latitude 71°n, and extends throughout north greenland. the park area includes the main 75 breeding area of the polar bear in greenland, and the greater part of the distribution area of the musk ox. access to the national park requires prior permission from the greenland authorities. 1974 hans meltofte, ornithological observations at kap tobin hans meltofte was employed at kap tobin weather station from march to september, during which period he made continuous ornithological observations. these included observations on journeys along the coast of southern liverpool land, and southwards across scoresby sund to kap brewster and steward ø (meltofte 1976). 1974 joint biological expedition to north-east greenland this large british expedition comprised two main groups, the ‘wader study group north-east green land expedition’ of 12 members led by g.h. green, and the ‘dundee university greenland expedition’ of 10 members led by j.j.d. greenwood. ornithological and zoological studies were made from base camps at holm bugt (traill ø), mestersvig, ørsted dal and antarctic havn (ferns & green 1975; ferns & mudge 1976; fletcher & webby 1977). the full report (green & greenwood 1978) gives positions of many of the numerous unofficial names used by this and earlier expeditions around mestersvig, in ørsted dal and around holm bugt in traill ø. [rgs and dpc report archives.] 1974 northern universities east greenland expedition: geoffrey halliday geoffrey halliday led a five-strong group that carried out botanical observations between mestersvig and fleming fjord. 1974 cambridge east greenland glaciological expedition: s.e. howarth a nine-member expedition led by s.e. howarth followed up the work of earlier cambridge university (uk) expeditions on roslin gletscher, where glaciological objectives included testing of a strain meter and thermal probe, and the surveying of stake lines. stakes were also surveyed on schuchert gletscher and arcturus gletscher. the expedition was supported by air drops by the british royal air force (raf). [rgs and dpc report archives.] 1974 ice king scientific expedition the ice-strengthened motor yacht ice king, com manded by michael tuson, sailed along the blosse ville kyst and into scoresby sund. a scientific party of botanists and geologists included r.m. sykes and s.r.a. kelly. geological studies were made in the kap leslie area of milne land (sykes & callomon 1979). 1974 sandhurst greenland expedition, ‘exercise snow goose’ a seven-man expedition from the royal military academy sandhurst, surrey, england, led by r.a.l. anderson, visited the bersærkerbræ region of the stauning alper. they camped at the junction of bersærkebræ and harlech gletscher, and were sup ported by a parachute drop of supplies by the british royal air force. due to poor weather only one ascent was made, of harlech on 16 august. 1974–79 grønlands geologiske undersøgelse (ggu) mapping projects in east greenland the geological survey of greenland (grønlands geo logiske undersøgelse, ggu) supported a number of general geological investigations in east greenland that made use of the base facilities at stordal set up for the ggu/aek radiometric investigations until 1977 (see above). some ggu groups operated independently, using the ggu cutter jytte based at mesters vig, or making use of chartered helicopters stationed at mestersvig airfield. 1974 – five mapping groups were active, and studies included a photographic reconnaissance of the blosseville kyst that clarified the distribution of the coast-parallel dyke swarm (watt 1975), sedimentological studies of mesozoic strata, and sampling for isotopic studies between 72° and 74°n (rex & gled hill 1981). 1975 – six groups were in the field, their projects including reconnaissance studies of the crystalline rocks between 72° and 74°n, studies of the triassic rocks on jameson land, and reconnaissance mapping of the blosseville kyst. 1976 – four groups carried out reconnaissance studies of the crystalline rocks between 72° and 74°n, 76 studies of basalts of hold with hope, and of triassic sediments. 1977 – five parties carried out a variety of studies, including work on lower palaeozoic rocks, tertiary basic rocks, and metamorphic studies in the crystal line complexes. 1978 – four groups continued work on projects including the basalts of the blosseville kyst and gauss halvø (upton et al. 1980), the crystalline complexes between 72° and 74°n (higgins et al. 1981), and on permian rocks on wegener halvø. 1979 – only one group was in the field, working on upper permian sediments west of schuchert flod. 1975 stirling university east greenland expedition: andrew ross andrew ross of stirling university (uk) flew to mesters vig and carried out studies in the vicinity and on boat trips to ella ø. more wide-ranging activities planned were frustrated by lack of transport. 1975 hans meltofte, ornithological observations around danmarkshavn hans meltofte was employed at danmarkshavn weath er station from march to september, and car ried out systematic ornithological observations (mel tofte 1977). 1975–76 ship-borne geophysical studies in the north atlantic geophysical investigations were carried out between jan mayen and east greenland using the rv explora for the bundesanstalt für geowissenschaften und rohstoffe, hannover, germany. one of the legs reached into the mouth of scoresby sund. the cruise continued in 1976 with further routes north and south of the mouth of scoresby sund. [dpc report archive.] 1975–76 geodætisk institut (geodetic institute) surveying 72°–76°n surveying was carried out between 72° and 76°n in 1975 by the danish geodetic institute from the motor cutter ole rømer with the aim of expanding a reliable triangulation network. in 1976 a number of the norwegian trigonometrical stations established by nsiu (lacmann 1937) were re-surveyed and syste m a tic gravity measurements undertaken. 1975–76 knud lauritzen, summer cruises knud lauritzen, owner of the danish j. lauritsen shipping company, visited many of the fjords between 72°–74°n in his small motor yacht saga dan in both 1975 and 1976 (fig. 23). soundings were made in sev eral small harbours and channels, including the sound between the head of alpefjord and dammen, and the narrow sound at strømnæs leading to the interior of röhss fjord. 1975–76 scottish scoresby land expeditions: e.a.m. walker 1975 – a four-man expedition led by e.a.m. walker carried out glaciological and botanical inves77 fig. 23. knud lauritzen, owner of the j. lauritzen shipping company, cruising in the east greenland fjords in 1976 with his motor-boat saga dan. tigations around oxford gletscher in the southern stauning alper (71°33´n), and also undertook a little climbing. the group had flown into mestersvig air field, and walked to gurreholm, from which a boat trip was made to scoresbysund to collect their equipment. many problems with transport were overcome, and the party was eventually flown to iceland from the rough airstrip in jættedal near scoresbysund. [dpc report archive.] 1976 – a party from edinburgh university (uk) led by e.a.m. walker carried out botanical and ecological work around bersærkerbræ and between kap peterséns and mestersvig. [dpc report archive.] 1976 karl herligkoffer climbing expedition karl herligkoffer led a seven-man german expedition to the spærregletscher region of scoresby land (72°n). ten ascents were made in the area around spærregletscher. 1976 austrian greenland expedition: helmut seerainer helmut seerainer led a six-man austrian expedition to east greenland, which made a number of climbs in liverpool land, and in the syltoppene in the northern stauning alper (hoff 1979). 1976 cambridge east greenland expedition: alan j. colvill a four-man cambridge univerity (uk) expedition led by a.j. colvill visited roslin gletscher (71°48´n) to undertake glaciological studies. these included measurement of a longitudinal profile and re-surveying of stake lines established by previous cambridge expeditions (see above ‘1968–75 east greenland ex peditions: keith j. miller’). [rgs report archive.] 1976 swedish-danish north-east greenland expedition a party of six made ornithological and quaternary geological investigations in the region of hochstetter forland, shannon, kuhn ø and sabine ø (74°30´– 75°30´n) between 26 may and 26 august (meltofte et al. 1981). the party was lifted into the area by helicopter, and subsequently made extensive journeys by ski and on foot. the main census area was in the vicinity of the nanok hunting station in southern hochstetter forland. two lakes (peters bugt sø and ailsa sø) were the subject of quaternary studies, and were named after nearby features. 1977 joint services expedition to liverpool land m.p.n. sessions led a 14-man joint services expedition (made up of members of the british armed forces) to liverpool land, with the object of making ornithological and botanical studies, and investigating hot springs. the group flew in at the end of may and sledged to carlsberg fjord (71°30´n). four peaks were climbed in northern liverpool land, including one unofficially named jubilee peak. [rgs report archive.] 1977 voyages of the rondø and santho two sailing ships visited the fjord region north of mestersvig in 1977, the rondø, a barque built by colin archer (the noted norwegian ship builder who built the fram), and the santho, a 32-foot pilot boat with a crew of four norwegians. both ships became trapped in the pack ice on the way home and were lost. the crews of the two boats were rescued by helicopter. 1977 university of dundee graduate expedition to north-east greenland: r.m.g. o’brien r.m.g. o’brien led a six-man group which travelled by sea to ella ø, and subsequently by inflatable boat in the kejser franz joseph fjord region. biological, geomorphological and hydrological studies were made, and the third ascent of petermann bjerg was made via knækdalen (rotovnik & søndergaard 1988). 1977 cambridge womens’ expedition to east greenland: v.m. haynes v.m. haynes led a seven-person group to roslin glet s cher (71°48´n), to continue the cambridge uni versity (uk) glaciological and geomorphological studies (see above ‘1968–75 east greenland expeditions: keith j. miller’ and ‘1976 cambridge east green land expedition: alan j. colvill’). they also cooperated with the ‘1977 cambridge east greenland glaciological expedition’ led by e.w. smith (see below). [rgs report archive.] 78 1977 cambridge east greenland glaciological expedition: e.w. smith e.w. smith led a four-man glaciological expedition from cambridge university (uk) to the roslin glet s cher area, where they co-operated with the ‘1977 cambridge womens’ expedition’. 1977 schwäbische grönland kundfahrt, stauning alper (german expedition to the stauning alper, east greenland): winfried baumgärtner this seven-person german expedition was led by winfried baumgärtner, and was lifted by helicopter to the borgbjerg gletscher area by helicopter. a total of 17 ascents were made around the head of the glacier, 16 of them first ascents (schloz 1979; rotovnik & søndergaard 1988). no names were given in the original report, the peaks being distinguished only by altitude, e.g. p. 2450. however, some names probably given by the expedition appear on the maps of the 1988 and 1992 scottish staunings expeditions. 1977–80 the british north-polar expedition: wally herbert wally herbert and allan gill had as their aim the first circumnavigation of greenland by dog sledge and umiak. they left thule in western north greenland in january 1978, but frustrated by very difficult ice conditions and non-availability of aircraft ‘restarted’ their journey from station nord in may 1978. by 14 june they had reached daneborg, and by 22 june loch fyne. they were airlifted to mestersvig airfield, used their umiak to retrace their steps to stordal, and then attempted to progress south of mestersvig. the attempt was abandoned in september 1978 due to difficult ice conditions, and a further attempt to resume their journey in 1979 was frustrated, again by ice (herbert 1979). further equipment was taken up in 1980, but a planned restart in 1981 was prevented by a telegraphists strike. 1978 army mountaineering association expedition, ‘exercise red eric i’ p.d. breadmore led a six-man expedition from the british army mountaineering association to the inner fjords of the scoresby sund region, climbing in the southernmost stauning alper, eastern renland, the bjørneøer and the inner part of vestfjord. trans port in the fjords was by inflatable boat, and included a circuit of milne land. one member injured an ankle in a fall, and was flown out to mestersvig airfield. sixteen peaks were climbed, of which 12 were reported as first ascents. some members made several long marches at the end of the expedition, including one from southern jameson land via hurry inlet to mestersvig. [rgs report archive.] 1978 familie journalen expedition a group of four led by jørgen bjerre, and financed by the danish magazine ‘familie journalen’, visited brøn lunds grav (79°09´n) on the 70th anniversary of the return of 1906–08 danmark-ekspeditionen. the jour ney to lambert land by helicopter was made in co-operation with the sirius relief and depot-laying flights. a memorial plaque was erected at the old de pot cairn. the body of brønlund had been ‘rediscovered’ in april 1963 by sirius, who had buried the remains beneath a large cairn and erected a brass plate (a gift from knud lauritzen) with the inscription ‘brønlund’s grav’. 1978–80 angus erskine ecological expeditions to north-east greenland angus bruce erskine [1928–2006] led a series of tourist expeditions to the mestersvig area, each with up to 13 participants. activities included walking tours, minor climbs, and ornithological and zoological observations. the expeditions continued in 1982 (see ‘1982–90 angus erskine ecological expeditions to north-east greenland’. [dpc report archive.] 1978–80 gi/ggu north greenland expeditions a three-year programme of geological and topographical surveying was carried out by a group from the geological survey of greenland (ggu) led by niels henriksen and members of the geodetic insti tute (gi), with a total of about 40 participants each year. the party was supported by three helicopters and a twin otter aircraft (peel & sønderholm 1991). in 1978 ground operations were in peary land, north of the region covered by this volume. super wide-angle aerial photography at a scale of 1:150 000 was carried out throughout north greenland, and also in east greenland from 76° to 82°n; the super 79 wide-angle aerial photography of northern east green land from 70° to 76°n was carried out in 1987. in 1979 both groups had limited activity in northern kronprins christian land, but the main activities were further north. a twin otter was lost in a refuelling accident at daneborg. the gi expedition was based at centrumsø in 1980, and control points were fixed over a large area of east greenland between 76° and 81°n. ggu field teams were active throughout kronprins christian land in 1980. 1979 ggu, ‘project eastmar’ ‘project eastmar’ was a grønlands geologiske under søgelse (ggu) energy research project, funded initially by the danish state through the energy agen cy, and subsequently included as part of a euro pean economic community supported geophysical project (‘project nad’ – see below). the eastmar project commenced in 1977, and in 1979 an aeromagnetic survey was flown over the continental margin off east greenland between 60°n and 80°n. survey operations were carried out by the western geo phys ical company of america, using a dc-3 aircraft operating out of narsarssuaq, kulusuk, mestersvig and reykjavik airports. a total of 63 000 line km of data were acquired (larsen & thorning 1980). 1979 ggu, ‘project dana’ 79 this project was part of grønlands geologiske under søgelse (ggu) geophysical investigations of the east greenland shelf, with special reference to its oil and gas potential. it was sponsored mainly as an energy-related research programme by the danish ministry of trade, industry and shipping, with support from the danish natural science research council. the survey was carried out by a ggu team of 10 using the dana, and a total of 10 000 line km of shallow seismic, magnetic and bathymetric profiles were completed between latitudes 60° and 71°30'n, including several lines within scoresby sund (larsen 1980). 1979 swedish north-east greenland expedition, myggbukta: magnus elander magnus elander and a companion made ornithological studies in the myggbukta region (73°29´n) of east greenland. this was a continuation of the studies by the ‘1976 swedish-danish north-east greenland expedition’ (elander & blomqvist 1986). 1979 zoological museum hurry inlet expedition j.m. hansen and n.o. jensen of the zoological mu seum, copenhagen, visited the head of hurry inlet (70°51´n) to carry out ornithological observations. between mid-may and early august they carried out intensive studies of waders. 1979 grea – groupe de recherches en écologie arctique (arctic ecology research group) a french expedition of four visited traill ø from june to august, and from a base camp at holm bugt car ried out ornithological and ecological observations. this was the first of a long series of grea expeditions to east greenland focused on the ecology of the arctic (grea 2003). [dpc report archive.] 1979 ggu / gdta, airborne remote sensing in east greenland grønlands geologiske undersøgelse (ggu) and the groupement pour le développement de la télé détec tion aérospatiale, toulouse (gdta), co-operated to carry out airborne remote sensing over se lect ed areas of east greenland between 70° and 74°n, with the support of the european economic community (eec) and the danish natural science research council. test areas with known types of mineralisation were overflown at different altitudes in august, using a boeing b-17 aircraft based at mestersvig air field. 1979 raoc greenland expedition, ‘exercise icy mountains v’ a nine-man british army expedition (royal army ordance corps) led by major a.j. muston visited the stauning alper and lyell land. from the head of dickson fjord (72°50´n) a traverse was made via agassiz dal and charpentier fjord to nordenskiöld gletscher, originally with the intention of making an ascent of petermann bjerg. this proved beyond the resources of the expedition, but ascents were made of several mountains in lyell land, including jeannet bjerg, argandhorn and snehætten (rotovnik & sønder gaard 1988). [rgs report archive.] 80 1980 dundee & milngavie north-east greenland expedition a six-member scottish group carried out systematic ornithological and botanical studies between skeldal and deltadal in the mestersvig area during august. their objectives were to expand the orithological studies of j.j.d. greenwood (university of dundee), and contribute to the botanical work of g. halliday (university of lancaster). [dpc report archive.] 1980 kaptajn ejnar mikkelsens minde ekspedition (captain ejnar mikkelsen memorial expedition): john andersen this two-man expedition led by john andersen made a journey from kap dalton (69°25´n) southwards to ammassalik by kayak, to commemorate the centennial of ejnar mikkelsen’s birth. the kayak journey was made in 59 days. their main mission was to search for traces of former inuit habitation (andersen 1980, 2005). [dpc report archive.] 1980 british army east greenland expedition, ‘exercise icy groove’ an expedition led by major h.w. beaves visited nathorst land and the stauning alper (rotovnik & søndergaard 1988). the party also gave some assistance to geoffrey halliday’s ‘british north-east green land expedition’. 1980 british army ‘exercise icy mountains vi’: a.j. muston a two-man british army group led by major a.j. muston formed part of the british north-east green land expedition (see below) which visited hoch stet ter forland. tours were initially made in the mesters vig region. subsequently climbs were made of wild spitze and matterhorn in the barth bjerge (75°30´n). 1980 british north-east greenland expedition: geoffrey halliday geoffrey halliday led a party to the bessels fjord and wollaston forland regions (74°20´–76°n), with the main purpose of carrying out ornithological and botanical studies. the group was landed by twin otter at the head of bessel fjord, where supplies had been dropped earlier by the raf (british royal air force). the party walked via langelv to mønstedhus, and then southwards. a move was then made to lindeman fjord by twin otter where further studies were carried out. [rgs and dpc report archives.] 1980 ‘ymer-80’ in the course of this swedish expedition with the icebreaker ymer to spitsbergen and the waters of northern greenland, observations were made along the coast of northern kronprins christian land. kilen was visited briefly by the geologist christian hjort (elg et al. 1981). [dpc report archive.] 1980–82 ggu, ‘project nad’ the objectives of ‘project nad’ involved geophysical mapping of the continental margin off east green land. the first part of the project, an aeromagnetic survey, was carried out as ‘project eastmar’ in 1979 (see above). the second part was a marine geo physical programme to collect seismic, gravity and magnetic data, scheduled to last from 1980 to 1982 and to cover the region from 69° to 77°n. the project was financed by the european economic community (eec) and the danish ministry of energy. the marine survey was contracted in 1980 to western geophysical company ltd. of america, who acquired 2610 line km of data using the survey vessel western arctic. in 1981 seismic profilers, oslo, were the contractors, and a further 2388 line km of data were acquired by the nina profiler. in 1982, western geophysical company ltd., again with the western arctic, completed the survey with 2794 line km of data (larsen 1983). 1981 geodætisk institut (geodetic institute) aerial photography super wide-angle aerial photography was carried out in east and south-east greenland between latitudes 62° and 70°n as part of a geodætisk institut project. 1981–82 italian stauning alper expedition: giuseppe dionisi giuseppe dionisi led eight-person groups from the italian alpine club to the stauning alper in both 1981 and 1982 (rotovnik & søndergaard 1988). in 1981 hjørnespids, norsketinde and dansketinde were climbed. in 1982 nine peaks around vikingebræ were climbed, again including hjørnespids, the first 81 ascent of norsketinde by the north ridge and the first traverse of dansketinde (dionisi 1983). 1981–88 gfm (grønlands fiskeriog miljøundersøgelser), gbu (grønlands botaniske undersøgelse), zoologisk museum, vildtbiologisk station kalø: jameson land activities a variety of environmental studies on the musk oxen, vegetation and birds of jameson land were carried out between 1981 and 1987, with particular reference to possible disturbances associated with intensive field activities during oil exploration. 1981 – vildtbiologisk station, kalø, on contract to grønlands fiskeriog miljøundersøgelser (gfm), began a survey of musk oxen in the spring, although a summer survey was suspended due to a telegraphists strike. 1982 – an aerial census of musk oxen was carried out by vildtbiologisk station, kalø, in april, and revealed a population of 3500–4000 animals in jame son land. follow-up studies on the ground were made between april and august, noting in particular the reaction of musk ox to helicopters. in july a total of 103 musk oxen were immobilised and tagged. also in 1982, the zoologisk museum, copenhagen, on contract to gfm, carried out ornithological studies, with particular reference to the goose population. grønlands tekniske organisation (gto) established an automatic weather station in central jameson land in august. 1983 – grønlands botaniske undersøgelse (gbu), on contract to gfm, carried out studies mainly of the distribution of vegetation types. vildtbiologisk sta tion, kalø, continued studies of musk oxen, a further 388 animals being immobilised and tagged. 1984 – groups from vildtbiologisk station, kalø, gbu and the university of copenhagen continued studies of musk oxen, vegetation and birds in the jameson land region. the zoologisk museum con tinued studies of the goose population, recording in excess of 6000 barnacle geese and 5500 pink-footed geese on the west coast of jameson land, ørsted dal and hurry inlet (madsen et al. 1984). gfm sponsored studies of the catch and distribution of marine mammals and seabirds utilised by the hunters of scoresbysund, including an aerial census of seals in kong oscar fjord and scoresby sund. 1985 – gfm continued studies of musk oxen, vegetation and birds in the jameson land area, and gbu also carried out studies of vegetation. the zoologisk museum made a special study of little auks on the coast of liverpool land and volquart boon kyst, recording a population of approximately 10 million birds. 1986 – studies were continued by gfm and gbu in jameson land, notably on the effects of human and helicopter disturbance on musk-oxen behaviour in the period january to march; during the summer vegetation studies were continued. 1987 – studies by gfm included an aerial census of barnacle geese (about 5000) and pink-footed geese (about 4000) in july and august. ground studies of breeding birds were carried out around gåseelv and ulveodde (inner hurry inlet). 1988 – studies of breeding birds were carried out around ugleelv and on the coast of hall bredning around jyllandselv (mortensen 2000). 1982 sheffield university north-east greenland expedition: bob andrews a seven-man group from sheffield university (uk) led by r.m. (bob) andrews visited the bersærkerbræ region (72°15´n), where geomorphological and gla cio logical studies were made, and a few peaks climbed. one party visited roslin gletscher to check supply depots left by earlier expeditions. [dpc & rgs report archives.] 1982 swedish north-east greenland expedition: magnus elander magnus elander led a two-man group to the myggbukta region (73°29´n) to continue ecological studies of birds, especially ducks and waders. 1982 grea – groupe de recherches en écologie arctique (arctic ecology research group) to east greenland christian kempf led a seven-person group on a follow-up to the grea 1979 expedition, visiting traill ø and vega sund. ornithological and ecological studies were continued (kempf 1986; grea 2003). 1982 skeldal expedition: keith j. miller keith j. miller led a four-man expedition to skeldal and the stauning alper, a follow-up to his previous 82 expeditions to the region (see ‘1968–75 east green land expeditions: keith j. miller’). 1982 stauning alper expedition: c.m. baker c.m. baker led a four-man climbing expedition to the region around the southern part of alpefjord. 1982 east greenland expedition: matti taponen a three-man expedition led by matti taponen visited the scoresby sund region. a helicopter search and rescue operation was carried out on 3 june after the expedition asked for assistance. 1982 ggu / gdta ground control studies the 1979 airborne remote sensing work carried out by grønlands geologiske undersøgelse (ggu) and the groupement pour le développement de la télé détection aérospatiale, toulouse (gdta) (see above, ‘1979 ggu / gdta, airborne remote sensing in east greenland’), was followed up by ground control studies in 1982 by a two-man group. areas in northern scoresby land, south-east traill ø and wegener halvø were visited. [geus archive.] 1982–83 nordisk mineselskab / atlantic richfield company (arco) under the terms of a non-exclusive exploration per mit, geophysical, geological and technical investigations were carried out in jameson land. exploration was continued in 1984 under the terms of an exclusive concession (see page 86 ‘1984–90 nordisk mine sel skab / arco oil exploration’). 1982–83 ggu, jameson land hydrocarbon studies grønlands geologiske undersøgelse (ggu) initiated a programme of source rock sampling, stratigraphical and sedimentological studies in jameson land, related to the planned oil prospecting of nordisk mine selskab and arco. in 1982 shallow drilling was carried out, and 265 m of core was obtained from 10 holes. in 1983 a further nine holes were drilled, each about 30 m deep (surlyk et al. 1984). 1982–83 henry dissing fungi expedition in 1982 a danish party of two led by henry dissing made studies of fungi in the mesters vig region and on ella ø. in 1983 the work was extended southwards to jameson land (dissing 1989). 1982–83 marathon oil company, wollaston forland a concession to prospect for oil and and gas in parts of wollaston forland was granted to marathon oil co., who undertook field work supported by a helicopter in the summer of 1982. in 1983 a group of geologists made a study of faulting in the jurassic and cretaceous sequence, led by finn surlyk. 1982–83 la croisiere glaces (crossing of the inland ice) christian gallissian had planned an expedition in 1981 to cross the inland ice from scoresby sund in east greenland to uummannaq in west greenland, from there continuing northwards, ultimately to reach the north pole. the 1981 plans were abandoned due to the telegraphists strike. in 1982 the expedition reached scoresby sund, but no further. in 1983 a renewed attempt met with success with a sledge crossing of the inland ice from scoresby sund to uummannaq. 1982–87 archaeological studies by grønlands landsmuseum in 1982 archaeological studies were made along the coast of jameson land between gurreholm and hurry inlet, the region in which oil exploration work was to commence in 1984. thirty-seven house ruins were registered, the interesting discoveries including a range of carved animal toys. in 1983 two winter houses were excavated, with amongst other things recovery of a large collection of ‘perle’, ornaments carved from bone and slate representing seals, birds and bears (sandell & sandell 1985). investigations were also carried out on the east side of hurry inlet. continued work in 1984 was concentrated on the west side of hurry inlet, where a new airfield (con stable pynt) to support oil exploration was to be con structed in 1985. 83 in 1985 a party of two carried out archaeological studies in the inland areas of jameson land. further activities in 1986 included ethnological and archaeological studies in western jameson land by a party of two, and preliminary excavations in the sydkap area by a six-man group, in cooperation with ilisimatusarfik (inuit institute). 1987 saw a continuation of excavations around syd kap, with reconnaissance activities in northern jameson land. 1982–90 angus erskine ecological expeditions to north-east greenland angus b. erskine continued his regular expeditions of l4–22 members to the mesters vig region, traill ø, hold with hope and hurry inlet, making walking and scrambling tours, and zoological and botanical observations. angus b. erskine had taken part in the british north greenland expedition in 1952–54, and also spent time in the antarctic during his career with the british navy. after his retirement in 1972 he founded his own small travel company ‘erskine expeditions’ that pioneered ‘ecotourism’ in the arctic with trips to svalbard, the canadian arctic and greenland. the company was taken over by ‘arcturus’ that continues to arrange small expeditions to east greenland and organises arctic cruises in east greenland waters. 1983 salford university mountaineering club greenland expedition: gerry mcculloch gerry mccullough led a nine-member climbing ex pedition from the university of salford (uk) to the stauning alper, setting up camp at the junction of bersærkerbræ and dunottar gletscher. eleven sum mits were climbed of which six were first ascents (rotovnik & søndergaard 1988). considerable time was spent filming an ascent for the bbc (british broadcasting company; peck 1984). 1983 expédition a.n.s., east greenland frédérik elin led a botanical expedition to scoresby sund on the 50th anniversary of the 1932–1933 international polar year expedition to scoresbysund. studies were made in the southern part of liverpool land. 1983 nathorst land reconnaissance expedition: j.l.w. walton j.l.w. walton led a party of nine to the furesø region of nathorst land (72°n). hydrographic surveys of dammen and furesø were made using inflatable boats. journeys included a walk via schaffhauserdal and violingletscher to the west end of furesø, and an ascent of sydgletscher. 1983 university of st. andrews expedition jean balfour and robert burton led a party of seven from the university of st. andrews (scotland) to the wollaston forland and sabine ø region, with climbing, botanical, ornithological and zoological objectives. 1983 dutch natural history expedition to north-east greenland: h.d. van bobemen h.d. van bohemen led a 12-member expedition to the mesters vig region between 23 july and 10 au gust, which made botanical and ornithological observations. [dpc report archive.] 1983 brathay trust north-east greenland expedition: steve f. newton a party of eight from the brathay trust led by steve f. newton made zoological investigations in ørsted dal and coloradodal (71°47´n), including a census of barnacle and pink-footed geese. the brathay charitable trust is based in the uk, and works with children and young people. it organises adventure training expeditions. [dpc & rgs report archives.] 1983 deutsche trans-grönland-expedition auf den spuren alfred wegeners (german trans-greenland expedition in the tracks of alfred wegener): arved fuchs an expedition led by arved fuchs retraced the steps of the 1933 alfred wegener expedition on the 50th anniversary of wegener’s death. after leaving mar mo rilik in west greenland on 8 may, the two-man party crossed the inland ice on skis, reaching east greenland at harefjord (70°55´n) on 15 july. as they were behind schedule the fjord ice had melted, and they were air-lifted to mestersvig by helicopter (fuchs 1984). 84 1983 french speleological expedition, ‘centrum 83’ j.-f. loubiere led a four-man group, supported by federation française de speleologie and societe arc tique française, to the centrumsø region of kron prins christian land (80°10´n). the french air force transported the group to station nord, from where they were air-lifted to centrumsø. a number of long foot traverses were made in the vicinity, including visits to the limestone caves of grottedalen, and caves south-west of centrumsø. the largest cave, at grottenfjeldet, has an opening nine metres high and was penetrated horizontally for 70 m (loubière 1989). [dpc report archive.] 1983 danish stauning alper expedition a danish three-man climbing expedition led by søren p. eisenhardt visited the stauning alper. their activities were restricted by bad weather to attempts on glamis borg and other minor summits (rotovnik & søndergaard 1988). 1983 robert peroni’s inland ice expedition an italian-german expedition of three led by robert peroni made an east to west crossing of the inland ice, beginning from ardencaple fjord (75°20´n) reached by helicopter on 17 june. they arrived at kraulshavn (74°07´n) in west greenland on 9 september (peroni 1992). 1983–84 k.g. swett geological expedition a four-man american geological expedition led by k.g. swett carried out studies of the upper pre cambrian and cambro-ordovician sequence in the fjord region of east greenland between 72° and 74°n. 1983–86 geodætisk institut (geodetic institute) activities the danish geodætisk institut carried out gravity measurements in jameson land and on the blosse ville kyst in 1983, as well as doppler-position determinations and gravity measurements north and west of mestersvig. in 1984 activities were concentrated in the south-west part of the scoresby sund fjord complex, and a network of fixed points was established across to the blosseville kyst. in 1985 super wide-angle aerial photography was flown over a large region from 70° to 76°n with an aircraft based at reykjavik. 1986 activities, that were co-ordinated with a grønlands geologiske undersøgelse (ggu) party, included a geodetic survey between ammas salik and scoresby sund. 1984 irish biological expedition to jameson land: david cabot an irish expedition of four members led by david cabot visited ørsted dal (71°47´n) in june–july with the main aim of studying the breeding ecology of bar nacle geese. a total of 644 barnacle geese and 8 pinkfooted geese were ringed. filming of their activities was released as an irish television film ‘valley of the geese’. [dpc report archive.] 1984 grea – groupe de recherches e écologie arctique (arctic ecology research group) east greenland expedition a french expedition of 14 led by christian kempf, sponsored by grea, continued their studies of 1979 and 1982, visiting geographical society ø and wolla ston forland to study tundra ecosystems (kempf 1986; grea 2003). [dpc report archive.] 1984 bedford college wildlife expedition: michael lea a four-member expedition from bedford college (lon don, uk) led by michael j. lea visited the mestersvig and scoresbysund areas to make wildlife sound recordings of birds. 1984 swiss geological expedition: christian böhm christian böhm of the university of bern, with the support of the schweizerische naturforschende gesellschaft, led a three-man expedition to the jame son land and mestersvig region to study stratabound lead-zinc-copper mineralisation. excur sions were made to the northern stauning alper, malmbjerg and south-west liverpool land. a minor summit on the north side of skjoldungebræ was climbed. 1984 italian climbing expedition to the stauning alper: sandro pucci sandro pucci led an italian climbing expedition of 85 eight persons to the stauning alper. despite bad weather, the expedition claimed 10 first ascents around gullygletscher that were all given italian names (anonymous 1985). 1984 jørgen brønlund mindeekspedition (jørgen brønlund memorial expedition): niels preben-andersen a 12-man danish expedition led by niels s. pre benandersen searched large areas of lambert land, kron prins christian land and danmark fjord (79°–81°n) for traces of the diaries and maps of l. mylius erichsen, n.p. høeg-hagen and jørgen brøn lund, the three men who died in 1907 during the 1906–08 danmark-ekspeditionen. no major discoveries were made. one man was evacuated by helicopter with a broken leg after falling 27 m down a crevasse on the glacier in nioghalvfjerdsfjord. [dpc report archive.] 1984 kayakekspedition station nord – scoresbysund john andersen and boas madsen made a journey by kayak and sledge along the coast of east greenland from station nord to scoresbysund. in the course of their voyage they shot six walrus and two polar bears, and were also rescued by helicopter from a position 55 km east of the norske øer after drifting out to sea in the pack ice. a second rescue operation was launched on 14 august after emergency signals were picked up by satellite, but this was a false alarm. they arrived at scoresbysund on 1 september (andersen 2005). [dpc report archive.] 1984 american geological expedition gerard c. bond and peter a. nickeson visited the region north of mestersvig in july. geological work including two weeks in the rarely visited nunatak region around eleonore sø (74°n), which was reach ed by helicopter. 1984 french stauning alper ski-mountain eering expedition: marc breuil a six-person french expedition led by marc breuil made a three-week, ski-mountaineering journey through the stauning alper between alpefjord and mestersvig in april–may. eight summits were climb ed, all over 2000 m (rotovnik & søndergaard 1988). 1984 swiss east greenland mountaineering expedition: alwin reither a group of four led by alwin reither made a twoweek mountaineering tour in the mestersvig region. 1984 österreichischer alpenverein (austrian alpine association) expedition: otmar resch otmar resch led a group of five on what was planned to be an ambitious mountain walking and skiing tour through the werner bjerge, across jameson land, and down the axis of southern liverpool land. the party arrived in june, but their activities were much hinder ed by melting snow. 1984–85 geological expedition to central east greenland a four-man expedition (m.j. hambrey, a.c.m. mon crieff, g. bylund and g. vidal) visited ella ø, ymer ø and suess land in 1984 to study precambrian tillites, as part of a north atlantic arctic synthesis. in 1985 studies were continued, and included visits to known tillite localities on charcot land and in paul stern land (moncrieff 1989; manby & hambrey 1989). 1984–90 nordisk mineselskab / arco oil exploration a consortium formed by atlantic richfield company (arco: 63.75%), arktisk minekompagni (a subsidi ary of nordisk mineselskab: 11.25%) and nunaoil (25%) was granted a 12-year concession to explore for and exploit oil and gas in a 10 000 square kilometre area centred on jameson land. the italian oil company agip took over half of arcos concession in the spring of 1988. a supply base was set up in hurry inlet in 1985 at constable pynt where a new airfield was built, and seismic surveys were begun in the winter of 1985–86. exploration drilling was initially planned for the summers of 1987 and 1988. seismic operations were suspended in the early spring of 1986 following a drastic fall in oil prices, but resumed in late 1987 after renewed negotiations led to a slightly revised concession. about 1500 km of seismic profiles had been acquired by the end of 1988, at a cost of 100 million dollars. in connection with the concession negotiations in 1984 the existing law governing 86 nordisk mineselskab and arktisk minekompagni and their concession rights was suspended. the con cession was given up without drilling in 1990. 1985 i.m. marsh college east greenland expedition: michael peckham michael peckham led a six-member expedition from i.m. marsh college, liverpool polytechnic (uk), to study the sedimentology and palaeoecology of raised marine sediments west of mestersvig from mid-july to early-september. climbs were made around ber særkerbræ, and 14 ascents were made, including the first ascent of d. eglin spids. [dpc & rgs report archives.] 1985 danish peary land expedition eigil knuth and henrik elling carried out archaeological studies in the region between frigg fjord (83°07´n) and lambert land (78°30´n), with helicopter support supplied by peter rutschman. 1985 ørsted dal botanical expedition to greenland geoffrey r. shaw led a four-man expedition to the ørsted dal region (71°47´n) with botanical objectives. several minor mountains were also climbed. 1985 ‘kilen 85’: eckart håkansson a party of six led by eckart håkansson carried out geological, botanical and ornithological studies in the kilen area of kronprins christian land. the party was flown in to kilen (81°12´n) from station nord by twin otter aircraft, and used all-terrain motorcycles for local transport (pedersen 1991). 1985 brathay trust east greenland expedition steve newton led a three-member party to the traill ø region for the brathay trust lasting from mid-june to early august. the brathay trust is a uk charitable organisation that arranges adventure holidays for young people. the ornithology studies, especially of geese, begun in 1983 were continued. a total of 117 geese were ringed, of which 85 were subsequently observed at the isle of islay, scotland, in november 1985. [rgs report archive.] 1985 gfm / ggu environmental studies environmental studies were carried out by grønlands fiskeriog miljøundersøelse (gfm) and grønlands geologiske undersøgelse (ggu) around mestersvig to investigate pollution arising from the mining activities of 1956–63. collections were made using the ship adolf jensen in august and september. 1985 belgium expedition to the stauning alper a group of eight climbers sailed with a ketch, via jan mayen, to east greenland. the ship was used as a base from an anchorage in dammen. climbs were made of dunottar bjerg and attilaborgen, and some members of the party explored the north shore of furesø and reached the col south-west of the head of the lake (borlée 1986). 1985 dundee university kejser franz joseph fjord expedition i.h.m. (ian) smart of dundee university (uk) and c.m.g. (malcolm) slesser were members of a fiveperson expedition which made botanical and orni thological studies in western frænkel land. the fourth ascent of petermann bjerg was made by both east and north-east ridges on 8 august. the southwest peak of trappebjerg was also climbed and named luxembourg spids, and one member of the party made a solo climb of gog (the fourth ascent). during their return to mestersvig an attempt was made on the highest peak of the syltoppene, but the party was repulsed by very poor rock (slesser 1987). 1985 geological excursion to east greenland claus heinberg and lars stemmerik led a party of 10 geologists from the norwegian oil company statoil to milne land, to make sedimentological studies. 1985–86 grea – groupe de recherches en écologie arctique (arctic ecology research group) to east greenland christian kempf led a party to ella ø in 1985 to under take ornithological and biological studies, a continuation of the 1978 activities by the groupe de recherches en écologie arctique (grea). studies were continued in 1986 in the region between mes 87 ters vig and myggbukta, with the aid of inflatable boats. [dpc report archive.] 1985–1987 expédition scientifique française au groenland est (french scientific expedition to east greenland) two small french expeditions visited the area around scoresbysund in 1985 and 1987, to carry out botanical and entomological studies. 1986 austrian alpine club, uk section, greenland expedition: john shrewsbury a group of seven led by john shrewsbury visited the inner scoresby sund region, landing by twin otter on eastern milne land. several peaks up to 1500 m high were climbed in july–august on the south side of charcot gletscher, and two 1200 m peaks west of bregnepynt (sales 1987a, b). [rgs report archive.] 1986 remote sensing studies on ymer ø a party of two led by john l. pedersen visited a test area on western ymer ø, to study the applicability of remote sensing techniques on landsat data in mineral exploration. 1986 expedition chamalieroise groenland (french expedition to greenland): bernard thomas a french expedition of nine led by bernard thomas visited strindberg land (73°50´n), making walking tours and climbs from a base near the mouth of brogetdal (rotovnik & søndergaard 1988). 1986 grønlands landsmuseum investigations at sydkap, scoresby sund hans kapel, henrik elling and tina møbjerg carried out archaeological excavations at a thule culture site at sydkap. 1986–88 ggu studies of the ‘onshore hydrocarbon potential’ in east greenland in 1986 a 19-member party from grønlands geo logiske undersøgelse (ggu) led by christian mar cussen and stefan piasecki worked out of a base camp at stordal, with the main activities on traill ø (72°30´n). source rock studies and shallow core drillings were undertaken in connection with oil exploration. in 1987 studies were continued with 17 participants, and extended northwards to kuhn ø (74°50´n) (marcussen et al. 1988). the ‘devonian basin project’ formed part of these studies and involved fieldwork in the period 1986 to 1988, supplemented by stereoscopic studies of vertical aerial photographs in ggu’s photogrammetric laboratory (larsen & olsen 1991). 1987 irish expedition to north-east greenland david cabot organised a three-person expedition to nordmarken (77°30´n), west of skærfjorden, lasting from end-may to mid-august. the principle aim was to study barnacle geese and pink-footed geese in their northern area of distribution. some helicopter assistance was provided by peter rutschman. numerous localities were given reference names in the expedition report (cabot et al. 1988), and a selection of them is included in this volume. [dpc & rgs report archives]. 1987 jørgen brønlund mindeekspedition (jørgen brønlund memorial expedition) niels s. preben-andersen followed up his 1984 expedition in search of traces of the lost members of the 1906–08 danmark-ekspeditionen. the five-man ex pedi tion visited southern kronprins christian land (79°45´n). 1987 british-danish palaeontological expedition to east greenland an expedition of five members led by svend e. bendix-almgreen, visited gauss halvø (73°26´n). extensive new collections of upper devonian tetra pods were made from around stensiö bjerg (bendixalmgreen et al. 1988). 1987 eric steen hansen lichen studies eric steen hansen carried out studies of lichen in the vicinity of scoresbysund, kap hope and kap tobin in july (hansen 1995). 88 1987 liverpool land expedition: michael lea michael and katherine lea, together with rob and sue david, visited the kalkdal area of liverpool land (70°50´n). [dpc report archive.] 1987: 2nd battalion royal green jackets greenland expedition: ‘exercise red eric ii’ a british army training expedition of eight members from the royal green jackets, led by robert a. chur cher, visited the inner scoresby sund region (70°–72°n) from mid-july to late september, using inflatable boats for transport. climbs were made of two 2000 m peaks north of stormpynt in the southernmost stauning alper, four 2000 m peaks in paul stern land north-east of arken, and a further two peaks in eastern paul stern land. the summits were all reported as easy, and none were given names. the return to constable pynt was made in extremely poor weather conditions. [dpc & rgs report archives.] 1987 geodætisk institut (geodetic institute) aerial photography the final season of the project to carry out super wide-angle aerial photography of all of greenland was completed in 1987, with coverage of the region 70° to 76°n in east greenland. opportunity was taken to fly supplementary routes to fill out gaps in the coverage of other regions. 1987 stauning alper expedition: françois wolf françois wolf led a party of six on a ski and climbing tour in the southern stauning alper in april and may. the party covered about 400 km on skis (rotovnik 1988). 1987 inland ice mass balance expedition a 13-man expedition from three german institutes undertook a largely airborne expedition to study the inland ice between ilulisat / jakobshavn in west green land and cecilia nunatak in east greenland. a helicopter visit to cecilia nunatak (72°30´n) was made on 10 july. geophysical flights were made using a specially equipped dornier research aircraft ‘polar 2’ from the alfred wegener institute, germany. 1987–88 renland glaciological expedition niels s. gundestrup (geophysical institute, univer sity of copenhagen) led a glaciological expedition of four members to the local ice cap on renland (71°15´n), as a prelude to drilling in 1988. field work was completed in seven days in early july, and an automatic weather station was erected for the meteoro logisk institut (danish meteorological institute). a 1988 follow-up expedition with six participants was carried out from 1 to 25 july, and an ice core drilled to bedrock at a depth of 325 m (johnsen et al. 1992). 1987–90 archaeological investigations on île de france (now qeqertaq prins henrik): eigil knuth in 1987 claus andreasen and henrik elling from grønlands landsmuseum joined up with eigil knuth and a student for an investigation of inuit sites in the dove bugt region that knuth had originally studied during his 1938–39 mørkefjord expedition. the group of four was provided with helicopter assistance by peter rutschman. sites at stormnæs, danmarks havn, rosio, rødeø and île de france were examined. the visit to île de france (now qeqertaq prins hen rik; 77°43´n) revealed more than 300 inde pendence ii ruins. in 1988, 1989 and 1990 eigil knuth returned to île de france with a few assistants and the support of the home rule authorities to continue his excavations, and the number of registered ruin sites rose to almost 500 (andreasen 2003). 1988 polarstern cruise ark-v/3 to scoresby sund region the r/v polarstern, research vessel of the alfred wegener institute for polar and marine studies (awi), sailed to the scoresby sund region in august and conducted a combined seaand land-based pro gramme. six land-based stations were established. reflection seismographic profiles were run across the passive continental margin and many geological samples were recovered. a further programme was carried out in 1990. 1988 british schools exploration society (bses) expedition to east greenland a 77-member bses expedition led by ray ward and george downy undertook a variety of scientific stud89 ies in the region south of mestersvig. this mountainering training expedition was carried out under the auspices of the duke of edinburgh award scheme and the royal geographical society, london. ski tours were undertaken on roslin gletscher, and oil drums and debris from mining operations were collected and tidied. [rgs report archive.] 1988 ‘exercise icy mountains viii’: milne land, greenland a nine-member british military expedition led by lt. col. a.j. muston visited milne land (70°43´n) in july– august. travel from constable pynt was by rubber boat. climbing and walking tours were undertaken, while the four civilian members of the expedition also made botanical collections. [rgs report archive.] 1988 scottish stauning alper expedition: john peden john s. peden led an eight-member expedition to the southern stauning alper. a planned ski traverse from sydkap in the south to kap peterséns in the north was frustrated by bad weather, and reached only as far as roslin gletscher. two new col crossings were made. [rgs report archive.] 1988 ‘exercise richmond circle’, first green howards greenland expedition: david charles johnson a party of eight from the british army regiment the first green howards, led by captain david charles johnson flew into constable pynt, with their boats to be used for transport. their journey from constable pynt to sydkap along the coast of jameson land was severely delayed by pack ice and bad weather. gurre holm was reached, but the party then abandoned their main objective of renland and retraced their steps; climbing and exploring was carried out in part of liverpool land. [dpc & rgs report archives.] 1988 bp wollaston forland a group from the bp (british petroleum) oil com pany visited wollaston forland (74°26´n) to make a geological training film. geological developments on shore east greenland are very similar to those in offshore areas of the north sea. 1988 wildfowl trust expedition to hold with hope a group from the wildfowl trust, a charitable con ser vation organisation based in the uk, sent a group to the hold wth hope area in july and august. in addition to their observations of birds, two wolves with two young cubs were observed; the female wolf was subsequently shot by another visiting group (turner & dennis 1989). 1988–89 harvard university palaeontological expeditions to east greenland: farish a. jenkins farish a. jenkins jr. of the museum of comparative zoology, harvard university led expeditions to the jameson land area in 1988 and 1989, with the pur pose of collecting vertebrate fossils from late triassic sediments. notable tetrapod samples were recovered, and spectacular footprint trails were observed (jenkins et al. 1994). the expeditions continued in 1991–92. 1988–90 grønlands geologiske undersøgelse (ggu) east greenland expedition this three-year grønlands geologiske undersøgelse (ggu) expedition led by niels henriksen operated from base camps on the west side of fligely fjord (1988) and at hvalrosodden (1989–90). geological objectives included systematic mapping of the region 75°–78°n (higgins 1994b; henriksen & higgins 2009). two helicopters and a twin otter aircraft provided transport to and within the area of research, and this logistical support was shared with other groups active in the region: the geodætisk institut (gi), alfred wegener institute for polar and marine studies (awi), grønlands landsmuseum, the zoo logical and botanical museums in copenhagen, and eigil knuth’s archaeological studies. 1988–90 greenland home rule government project: biologisk–arkæologisk kortlægning af grønlands østkyst mellem 75°n og 79°30´n (biological–archaeological mapping of east greenland between 75° and 79°30´n) these investigations, sponsored by the greenland national museum in nuuk, were a co-operative ven ture between the zoological and botanical museums 90 in copenhagen and the greenland na tional museum, and were focused on the north-east greenland national park. in 1988 botanical and zoological observations from the air (72°–78°n) were made to select the areas for 1989 ground observations. the 1989 botanical, ornithological and entomological studies were carried out at 14 localities between bessel fjord and zachariae isstrøm (76°–78°30´n), with particular reference to areas with breeding geese and other birds and the distribution of musk oxen (boertmann et al. 1991). botanical studies continued in 1990 (boertmann & forch hammer 1991). archae ological studies covering the entire area were carried out in 1989 and 1990 by zodiac rubber boat and helicopter by claus andreasen and henrik elling. helicopter transport in 1989 and 1990 was supplied by arrangement with the ggu base camp at hval ros odden. the ‘1989 danmarks radio, nordøst grønland rejse’ (danish broadcasting corporation, north-east greenland visit) (see be low) was an activity under this project. 1988–1991 grea – groupe de recherches en écologie arctique (arctic ecology research group) to north-east greenland: benoît sittler a nine-person expedition led by benoît sittler visited the karupelv region of traill ø in 1988. this was a follow-up of earlier grea expeditions. a similar expedition in 1989 visited the same region. in 1990 there were two groups, a group of four led by benoît sittler based at karupelv on traill ø, and a second group of four led by christian kempf engaged in ornithological studies between kong oscar fjord and myggbukta. in 1991 benoît sittler continued studies of birds, animals and snow-melt patterns in the karupelv area of traill ø with a nine-person party (grea 2003). [dpc report archive.] 1989 newcastle university east greenland expedition: s.j. munro s.j. munro was leader of a group of six persons from newcastle university working in the mestersvig re gion in july–august on a study of the impact of the former lead mine on sediments and the floral com munities. four members made climbs in the stauning alper. [dpc & rgs report archives.] 1989 danmarks radio, nordøstgrønland rejse (danish broadcasting corporation, north east greenland visit) under the 1988–90 greenland home rule govern ment project (see above) the danish broadcasting corporation (danmarks radio: dr) made a summer visit to northern east greenland to make a series of television films of the activities carried out under the 1988–90 greenland home rule government’s pro ject: biologisk-arkæologisk kortlægning af grøn lands østkyst mellem 75°n og 79°30´n; they also visited the sirius headquarters at daneborg, the danmarkshavn weather station, the ggu geological activities based at hvalrosodden (see above), and eigil knuth’s archaeological excavations on île de france (now qeqertaq prins henrik). television pro grammes were later broadcast on dr television and in a number of countries. 1989 nordøstgrønlandsekspeditionen (north-east greenland expedition): jan juel-brockdorff jan juel-brockdorff and a companion visited the area of nordostrundingen (81°15´n) in a continued search for traces of the lost records of the 1906–08 danmarkekspeditionen. bad weather in mid-august led to their evacuation by helicopter to station nord. [dpc report archive. 1989 mylius-erichsens mindeekspedition (mylius-erichsen memorial expedition): finn rasmussen a six-man group led by finn rasmussen working from a base at marmorvigen (80°05´n), followed the coasts of holm land and hovgaard ø in a continued search for traces of the lost records of the 1906–08 danmark-ekspeditionen. [dpc report archive.] 1989 greenland milne land expedition: malcolm sales malcolm sales led a group of eight on a climbing expedition to milne land in august, landing by twin otter at the rough airstrip on the coast between breg ne pynt and charcot havn. seven peaks up to 2080 m high were climbed around the glacier-filled valley korridoren. [dpc & rgs report archives.] 91 1989 ‘exercise snow dance’, british army expedition to liverpool land: m.t. king an eight-member british army expedition led by major m.t. king took part in adventure training and exploration in southern liverpool land. nine peaks were climbed, including korsbjerg, where a cairn with a record from 1933 was found. [dpc & rgs report archives.] 1989–90 hvalrosundersøgelser i nordøstgrønland (walrus studies in north-east greenland): erik born erik w. born and lars ø. knutsen undertook studies of walrus at a haul-out location at lille snenæs on the south coast of germania land in august 1989 and 1990. their observations indicate that about 52 male walruses used the lille snenæs site in 1990, with the maximum number of walruses on a single occasion numbering 48. a total of 12 walruses were equipped with satellite radio transmitters, so that their wanderings, diving frequency and swimming speed could be recorded for up to six months (born & knutsen 1991). 1989–90 alfred wegener institute, east greenland expedition three participants from the alfred wegener institute for polar and marine studies (awi) made glaciological studies at the ice-sheet margin on storstrømmen (77°n) and west of dronning louise land in july and august. studies were part of an eu-supported project ‘climate change on a century time scale’ and included the present dynamic and climatic conditions, fluctuations of the position of the glacier and velocity variations (reeh et al. 1994; weidick et al. 1996). lo gistics were shared with the grønlands geologiske undersøgelse (ggu) expedition based at hvalros odden. [geus archive.] 1990 british north-east greenland expedition: geoffrey halliday geoffrey halliday led a botanical expedition from the university of lancaster to the kuhn ø and wollaston forland areas (74°–75°n). a total of 18 persons divided into two groups were involved in botanical, ornithological and faunal surveys. one group undertook an archaeological survey of kuhn ø and eastern th. thomsen land, with a visit to mågenæs in grandjean fjord. [dpc report archive.] 1990–1992 ponam (polar north atlantic margins) project the ponam project was a study of the late ceno zoic climatic and environmental history of the european arctic, focusing on the last inter glacial/ glacial cycle. in east greenland the main groundbased work was in 1990 on jameson land and in 1992 on hochstetter forland and wollaston foreland. the almost 50 participants in the ponam project were mainly from the scandinavian countries, germany and the uk (funder et al. 1994). in 1990 the investigations on land were complemented by a marine geological survey of scoresby sund by the polar stern (see below). 1990 the polarstern geophysical cruise ark-vii/3 in scoresby sund the polarstern, research vessel of the alfred wege ner institute for polar and marine studies (awi), bremen, carried out a programme of geophysical work in scoresby sund and on the adjacent shelf in september. these marine investigations of the late quaternary sedimentary record were a supplement to the onshore studies of the ponam project (dowdes well et al. 1994). 1990 british schools exploring society (bses) expedition – mestersvig region ray ward again visited the mestersvig region with a party of 48 young people as a follow-up of the 1988 expedition organised by the british schools ex ploring society (bses); see above. tours in the vicinity of mestersvig were extended to deltadal, schu chert dal and roslin gletscher, where activities included climbing, and biological and glaciological studies. departure from mestersvig airfield was de layed by a week due to heavy rain that closed the runway. 1990 hold with hope insect project a finnish group of two led by erkki m. laasonen visited the hold with hope region (73°45´n) to study the insect fauna, part of a circum-arctic project. 1990 bristol university north-east greenland expedition: jonathan rowe jonathan rowe led a six-member party to the northern stauning alper in july and august. activities 92 included investigations of meltwater streams on glaciers and studies of atmospheric pollutants. climbs were made on beaumaris, tintagel and spiret (ber zaerkerspire). [dpc & rgs report archives.] 1990 casp east greenland project: chrispin day the cambridge arctic shelf programme (casp, uk) made studies of devonian sedimentation and tectonics in the kong oscar fjord and kejser franz joseph fjord region with a party of four led by chrispin day. 1990 ‘exercise green ice’, royal military college of science: andrew b. syme a british expedition of eight members from the british royal military college of science, shriven ham, led by andrew b. syme visited the stauning alper in july and august. activities included glaciological and meteorological observations, skiing and climbing. on 7 august an ascent was made of an 1800 m high peak on the south-west side of schuchert gletscher, which they named mt. shrivenham. [dpc & rgs report archives.] 1990 icelandic greenland expedition: ingimundur stefansson ingimundur stefansson led a five-person expedition from the icelandic alpine club to the stauning alper. the group flew into mestersvig on 21 july, with an air-drop of equipment on bersærkerbræ on the way. from mestersvig the party walked in to bersærker bræ, and over the next five weeks climbed tintagel fjeld, kensington, spiret, dunottar bjerg and black wall. a failed attempt was made on glamis borg (rotovnik 1991). 1990 expedition greenland: bernard thomas a nine-person group led by bernard thomas attempted to reach petermann bjerg via knækdalen, to which they had been transported by the icebreaker cari boo. the party attempted to ascend knækdalen on the west bank of knækelven, which was in flood and proved uncrossable. the ascent was given up. the party was picked up by members of the french grea expedition, and after a difficult return journey by rubber boat to ymer ø, was flown to iceland by twin otter. [dpc report archive.] 1991 british schools exploration society (bses) north-east greenland expedition: dave walker a party of 80 persons, made up of 17 leaders and 63 young ‘expeditioners’, under the overall leadership of dave walker, visited the mestersvig region in july and august. an integrated programme of scientific studies, adventure and personal challenge was undertaken in the skeldal and deltadal areas of scoresby land. [dpc & rgs report archives.] 1991 scottish stauning alps expedition: alex erskin alex erskin led a four-man climbing expedition to the bersærkerbræ area of the northern stauning alper. ascents were made of beaumaris, elizabeths minde, spiret and pimlico. a summit north of tintagel was climbed and called bear peak. [dpc report archive.] 1991 nordøstgrønlands ekspedition (north-east greenland expedition): jan juel-brockdorff jan juel-brockdorf with one companion undertook a thorough search of the coast of lambert land and the islands in front of nioghalvfjerdsfjord, for traces of the lost members of the 1906–08 danmark-eks pedi tionen between may and august. [dpc report archive.] 1991–1992 vertebrate palaeontological expedition to jameson land: farish a. jenkins a follow up of the 1988–1989 expeditions, farish a. jenkins jr. led an 11-strong expedition to the ørsted dal – allday dal region of jameson land in july– august 1991. excellent collections of triassic vertebrate fossils were made (jenkins et al. 1994). a further eight person expedition in july–august 1992 was somewhat hindered by extensive snow cover. [dpc report archive.] 1991–1998 greenland wolf research project the danish scientist ulf marquard-petersen began a long-running research project on the ecology of arctic wolves in greenland in 1991. fieldwork was carried out in nansen land (north greenland, 83°n, 1991), hold with hope (74°n, 1992–1994), peary 93 land (north greenland, 83°n, 1995), wollaston forland and hold with hope (74°n, 1996), liverpool land (70°30´n, 1997) and germania land (77°n, 1997), kronprins christian land (80°n, 1998) and hold with hope (74°n, 1998). numerous observations of wolves, some with recent young, have been documented (marquard-petersen 1994). [dpc re port archive.] 1991–2003 grea/cedme east greenland expeditions the groupe de recherches en écologie arctique (grea – arctic ecology research group) continued their studies in east greenland in 1991, and from 1992 another long-term project was launched by the centres d’études et de documentation sur les milieux polaires (cedme) (grea 2003). 1991 – grea again based their activities at karupelv with a group of six persons. 1992 – grea continued their routine monitoring studies around karupelv with a six-person group, and cedme undertook botanical, ornithological and mammal studies in the fjord region 72°–74°n with a three-person party. 1993 – grea started their karupelv studies very early this year, in may, and the seven members experienced problems with a polar bear that destroyed two tents. 1994 – the grea group of eight was joined in 1994 by a danish polar center participant on their monitoring studies around karupelv. they were witness to a peak in the lemming population. 1995 – a grea group of seven again undertook routine monitoring studies around karupelv, and a cedme group of four worked in the fjord region 72°–74°n. 1996–2000 the grea groups of four to eight members continued their monitoring studies around karupelv. the cedme group worked mainly in the fjord region 72°–74°n in 1998–2000, on a project mainly focused on lemming predators. in 1999 radiocollars were fitted to 17 lemmings to track distances covered (grea 2003). 2001–03 the grea monitoring project based around karupelv was continued with groups of up to five persons. in 2002 an additional group of five undertook studies around kejser franz joseph fjord using kayaks for transport. a cedme four-person group continued their studies of lemming predators in the fjord region 72°–74°n (grea 2003). see also ‘2003 ecopolaris (grea) expedition’, ‘2004 ecopolaris (grea) tara 5 expedition’ and ‘2007 grea sagax-revo and ecopolaris expeditions’ below. 1991–present: nanok expeditions all danish and norwegian hunting stations and hunting huts within the national park were granted preservation status in 1987. after trapping ceased in 1952, the company continued to exist with j.g. jen nov as director until 1976, when he was succeeded by mogens graee. the old ‘østgrønlandsk fangst kompagni nanok’ was liquidated in november 1990 and the assets passed over to mogens graee. in july 1991 six enthusiasts with interests in northern east greenland met at graee’s cottage in jutland, and this was followed by a two-man summer expedition in august–setember 1991, and a few months later the rebirth of nanok. on 12 january 1992 the official name of the company was changed to ‘nordøst grønlandsk kompagni nanok’ (north-east green land company nanok), usually known as ‘nanok’. the vision of the new company was to: “disseminate knowledge of north-east greenland and its cultural history, and to contribute to maintenance of the cultural relics and buildings of the area...” (p.s. mikkelsen 2008, p. 47). from 1991 members and associates of nanok began a regular programme of repairs and maintenance with between three and 10 persons in volved each summer. 1991–1992: during the first two years of the programme, repairs were carried out on the zacken berg hunting station. 1993: a three-man group restored the loch fyne station and arvehytten in july–august. 1994: a two-person party undertook maintenance in july–august of the sandodden and moskusheimen hunting stations. 1995: a three-member party renovated the hunting station of ny jónsbu in the ardencaple fjord region in july–august. 1996: a group of four persons undertook repairs of hochstetter station, known also under the approved name nanok. 1997: the norwegian hunting stations at kap humboldt (known as humboldt), and at kap peter séns were renovated. 1998: repairs were continued at kap peterséns and the hut adjacent to the burnt-down eskimonæs sta tion was restored. 94 1999: maintenance activities were concentrated on the danish germaniahavn station on sabine ø, and the norwegian stations hoelsbo in moskus oksefjord and myggbukta in mackenzie bugt. 2000: renovation of hoelsbo was completed, and repairs were made of the hut at kap ovibos. a cultural–historical collection of artifacts relating to the hunting period in east greenland was established in a building (‘hotel karina’) at sandodden. 2001: the station at antarctic havn was restored, and repairs carried out on the huts at kongeborgen and holm bugt. 2002: restoration of the myggbukta station, begun in 1999, was completed, and extensive repairs made of herschellhus on wollaston forland. the varghytta in blomsterbukten was rebuilt. 2003: two groups undertook a major programme of registration of the status and exact (gps) positions of huts and stations between 72° and 75°n. 2004: registration of the condition and positions of huts was continued, again in two groups, reaching as far north as hochstetter forland (75°25´n). 2005: two groups continued the status pro gramme of registration and photographing of hunting huts, and fixing of their positions. 2006: two groups continued registration of the status of huts. the northern group also gave special attention to the preservation of villaen / danmarks minde at danmarkshavn. 2007: the programme of registration and repair of huts was continued, with particular attention given to the condition of many huts previously repaired in the period 1991–2002. 2008: two groups continued the programme of restoration of huts. a southern team restored and re paired the huts maristua, arentzhytten, bjørnheimen and noa sø hytten. a northern team repaired the huts elvsborg, fiskerhytten, bjørnnesstua and leirvågen. [dpc report archive.] 1992 scottish bersærkertinde expedition greenland: stan pearson a four-member scottish climbing expedition led by stan pearson visited the northern stauning alper in june–july, approached from alpefjord and sefström gletscher. deep fresh snow and high meltwater streams proved considerable hinderances. failed at tempts were made on bersærkertinde and attila borgen. [dpc & rgs report archives.] 1992 scottish staunings expedition: john peden a party of six scottish and french climbers made a splendid 18-day south-to-north, ski traverse of the stauning alper in may. dropped off by twin otter on the sea ice of nordvestfjord near stormpynt, the tra verse began on 7 may with an ascent of oxford glet scher. a total of eight passes were crossed and three sum mits climbed, ending with a descent of skjold ungebræ to reach kap peterséns on 24 may (peden 1993). [dpc & rgs report archives.] 1992 eclogite expedition to danmarkshavn: jane a. gilotti jane a. gilotti continued her studies of eclogites, be gun during the 1988–90 ggu east greenland expeditions, concentrating her efforts in the vicinity of danmarkshavn in july. [dpc report archive.] 1992 scottish mountaineering club expedition to the stauning alps: w. wallace a nine-person group led by w. wallace visited the stauning alper in may. they were assisted by an airdrop of equipment and provisions near gefion pas. the group split into two parties. one party climbed harlech fjeld and minor peaks around blyklippen. the second party climbed dunottar bjerg and beau maris fjeld. other peaks were attempted, but the climbing parties were repulsed by snow conditions. [dpc report archive.] 1992 ‘high latitude astronomers expedition’ to east greenland this seven-person climbing party comprised two british, four canadian and one norwegian climber. they flew into mestersvig on 24 july, and reached bersærkerbræ via skelbræ, kishmul gletscher and glamis pas. they were frustrated in many of their objectives by poor weather and difficult snow conditions. ascents were made of richmond and harlech fjeld (aarseth 1993). they flew out from mestersvig on 8 august. [dpc report archive.] 95 1992 dr–derude til nordøstgrønland (danish broadcasting corporation outsidebroadcast unit visits north-east greenland): mogens guldbrandsen mogens n. ‘gulli’ guldbrandsen, for many years leader of the sirius sledge patrol, visited northern east greenland between mid-february and mid-may. the group included a two-person tv film crew, a former sirius patrol member ‘tavse’, two sledges and 22 dogs. the party flew from iceland to mestersvig, and after about 10 days of preparation and training sledged from mestersvig to daneborg (74°18´n), with periodic stops at various former trapping sta tions for filming purposes. after a five-day stopover at daneborg the group was flown on 23 april by twin otter to kap stop (76°38´n), with a short stop at alabamahuset on shan non on the way. from kap stop the sledge journey was continued along the west side of dove bugt and the south coast of germania land to danmarks havn (76°46´n). a 12-day stopover at danmarkshavn was followed by a further twin otter flight on 18 may, via brøn lunds grav, to station nord and kap morris jesup, the north point of greenland. the group was later flown back to mestersvig, where they arrived on 22 may. as a result of these activities, a series of excellent short film episodes were broadcast on danish television in 1993. [dpc report archive.] 1992–2000 british north-east greenland project rob david organised a series of expeditions, surveying archaeological sites, with subsidiary botanical and ornithological observations. some notes on the archaeological observations were published by david (1995, 1999). michael j. lea also organised and led many of the expeditions. [dpc report archive.] 1992 – a seven strong group visited the clavering ø region in july–august. 1993 – an eight person group visited the lyell land region, documenting archaeological sites on hammer ø, kap lagerberg and at kap harry on ella ø. 1994 – a six member group visited the strindberg land region, studying archaeological sites at primula bugt, nordfjord and kap ovibos. botanical studies were concentrated in brogetdal. 1995 – seven persons visited the bjørneøer region of inner scoresby sund. study areas extended from northern milne land, through the islands of the bjørne øer to sydkap. 1996 – a second visit was made to milne land in the scoresby sund region. 1997 – kejser franz joseph fjord was visited by a nine-person group led by michael lea and an attempt made on petermann bjerg, turning back 300 m from the summit due to dangerous ice conditions. 1998 – rob david led an eight-person group to southern clavering ø. investigations were mainly bo tanical. 1999 – michael j. lea led a group to the region around the southern coast of clavering ø, making walking tours and wildlife studies in july–august. observations of walrus were reported. 2000 – mountaineering, botanical and wildlife observations were made on clavering ø and vicinity, led by michael j. lea. 2005–2008 – see ‘british north-east greenland project’ below. 1993 mylius-erichsens mindeekspedition (mylius-erichsen memorial expedition): finn rasmussen a danish four-man group continued the regular expeditions looking for traces of the lost records of the missing members of 1906–08 danmark-ekspedi tio nen. in july–august areas were visited on the west side of danmark fjord, along skjoldungeelven, and in southern kronprins christian land between blåsø and kap bernhoft. no significant new relics were dis covered. [dpc report archive.] 1993–1995 ggu/geus lambert land and kronprins christian land mapping project a three-year regional geological mapping project led by niels henriksen was commenced in 1993 by ggu, aimed at production of a map sheet in the survey’s 1:500 000 scale series (sheet 9: lambert land) (fig. 24). the lambert land map sheet covers the region between jökelbugten (78°n) and northern kronprins christian land (81°n). in all three years work was car ried out from a base camp at the west end of centrum sø, with the field parties supported by 1–2 helicopters and a twin otter aircraft (henriksen 1996; fig. 25). in 1995 grønlands geologiske under søgelse (geological survey of greenland: ggu) was merged with danmarks geologiske undersøgelse (geo logical survey of denmark: dgu) to form a new insti tute, the geological survey of denmark and green land 96 97 (geus), and this and other geological projects were continued as geus projects (see ghisler 1996). collaboration was carried out with two german geoscientific institutes, the alfred wegener institute for polar and marine research (awi – bremerhaven) and the federal institute for geosciences and natural resources (bgr – hannover). 1993–1995 glaciological research in northern east greenland scientists from the alfred wegener institute for polar and marine research (awi) continued their studies of 1989–90 on storstrømmen, setting up and measuring stake lines to determine velocities and establishing a number of automatic climate stations. in 1995 similar studies were made on the margin of the inland ice south-west of centrumsø (henriksen 1996). 1993–2003 ggu/geus/dlc east greenland field activities from 1993 onwards summaries of activities in green land were issued each year (‘feltaktiviteter i grøn land’) covering the work planned to be carried out by ggu/geus/dlc (grønlands geologiske under søgelse – ggu / danmarks og grønlands geologiske undersøgelse – geus / danish lithosphere center – dlc). the activities planned in northern east green land, extracted from these summaries, are given below. 1993: the main ggu-sponsored activities included the first summer of the 1993–1995 regional geo logical mapping programme (see separate entry above). another large group initiated a planned three-year programme of studies of post-caledonian sedimentary basins with sequence stratigraphic studies in jame son land. a related project on the onshore hydro carbon potential of east greenland continued with sampling of lower cretaceous sequences for dinoflagellate cysts. ablation-climate studies were carried out on the margin of the inland ice near kronprins christian land, and of glacier outlet dynamics on storstrømmen, the glacier at the north-east margin of dronning louise land. these were part of joint projects between ggu, the alfred wegener institute for polar and marine research (awi) and other institutes, related to global climate change. 1994: as part of the ggu/dlc ‘east greenland volcanic rifted margin project’, systematic stereophotography was undertaken in august 1994 of the lava plateau basalts and the coastal dyke swarm between 66° and 70°n using a twin otter aircraft. a total of 1600 km of mountain sides were photo graphed. this was part of the dlc (danish litho sphere centre) programme of studies on the opening of the north atlantic ocean. 1995: the project ‘resources of the sedimentary basins in north and east greenland’ was a joint project that involved geologists from geus (formed in 1995 by a merger of the geological survey of green land – ggu, and the geological survey of den mark – dgu; ghisler 1996), the universities of copenhagen and aarhus and the danish environ mental research institute (dmu). work in 1995 was carried out in the fig. 24. segment of the 1:500 000 geological map lambert land published by the geological survey of denmark and greenland (geus) after the 1993–1995 mapping project ( jepsen 2000). the segment extends from the west end of centrumsø, where the base camp was located, eastwards to lynn ø and dimphna sund. franklinian basin of north green land and in the east greenland rift basins (traill ø). studies of the petroleum systems in the wandel sea basin in kronprins christian land were co-ordinated with the ggu/ geus lambert land and kronprins christian land mapping project. a large international field team of 34 scientists carried out fieldwork south of scoresby sund (70°n) as part a continuation of the ggu/dlc (now geus/dlc) project on the east greenland volcanic rifted margin. two helicopters allowed ac cess to areas previously considered inaccessible. 1996: the project ‘resources of the sedimentary basins in north and east greenland’ was continued with nine field teams active between 71° and 74°n. investigations of the pleistocene sedimentary record of the falsterelv area of jameson land were carried out, a continuation of earlier studies in the same area. a programme of glaciological research was initiated on the nioghalvfjerdsfjorden glacier, an international project involving geus and the danish polar center as danish partners; this was part of a threeyear project studying ice-sheet response to climate change. 1997: the geus regional mapping project of the kong oscar fjord region (1997–1998) is described separately below. petroleum-geological activities were continued by five field teams working particularly on permian to cretaceous sedimentary successions. continued glaciological studies around the nioghalvfjerdsfjorden glacier revealed several pingos and pingo-like structures, the northernmost known in greenland. geus carried out an airborne electromagnetic and magnetic survey over northern jame son land (part of project ‘aem greenland’; stemp 1998; rasmussen et al. 2001). 1998: continued quaternary studies around the ni oghalvfjerdfjorden glacier included visits to søndre mellemland and île de france / qeqertaq prins henrik. the main geus activities were the second and final summer of the kong oscar fjord regional mapping project (see below). 1999: the main geus activity in east greenland was a visit to southern renland (71°n), to determine the relationships between previously described oro genic deformation and c. 935 ma magmatic activity. samples were collected for isotopic age determinations (leslie & nutman 2003). 2000: studies of vendian–ordovician stratigraphy were conducted on ella ø in association with geologists from the geological museum, copen hagen. in the mestersvig area activities included an assessment of the changes in periglacial processes since the studies by a.l. (linc) washburn in 1955– 64. airborne hyperspectral data were acquired over selected areas of northern east greenland be tween 71°30´ and 73°30´n, part of a collaborative ven ture by nine european research organisations and two min ing companies (projects ‘mineo’ and ‘hyper green’). 2001: activities in northern east greenland in cluded sample collection from known ore showings on clavering ø (project ‘hypergreen’); fossil collections from the cambrian successions on ella ø and albert heim bjerge; and investigations of the thin sediments at the tops of lava flows in the kap dalton 98 fig. 25. the ggu/geus base camp at centrumsø, kronprins christian land. two small helicopters transport two-person geological field teams to new camp sites at about 6–7 day intervals. the twin otter aircraft was mainly used for transport of helicopter fuel from station nord to centrumsø. the large tent holds supplies for the aircraft, while the small tents at left house base camp staff and geologists. photo: jakob lautrup. area and the northern blosseville kyst, with in particular a search for oil seeps. 2002: geus was responsible for leading a field excursion in jameson land and milne land for a number of oil companies interested in comparisons with the jurassic–cretaceous strata offshore the nor we gian coast. 2003: investigations were concentrated in the scoresby sund area, with particular reference to the cretaceous–tertiary sediments underlying the ter tiary basalt succession at bopladsdalen west of kap brewster. [geus archive.] 2003–present: see ‘2003–present geocenter den mark below. 1994 mountain adventure kayak expedition an eight-person group undertook kayak tours ex tend ing from revet west of clavering ø, to dane borg, eskimonæs and loch fyne. 1994 cardiff university greenland expedition: gary timms gary timms led a six-strong party from cardiff univerity (uk) to the bersærkerbræ area of the stau ning alper in july–august. studies were made of temperature and pore-water pressure on periglacial slopes. climbs were made on dunottar bjerg and glamis borg. [rgs report archive.] 1995 botanical studies in kronprins christian land six botanists from the university of copenhagen and the university of münster, divided into three twoperson teams, studied plants, lichens and mosses in kronprins christian land during july and august. they were moved periodically by the helicopters at the geus base camp at centrumsø (henriksen 1996). 1995–2007 arild andresen caledonian geological studies east greenland arild andresen, of the university of oslo, began a major project in 1995 to study aspects of the collisional and extensional history of the caledonides, and the post-caledonian sedimentation linked to orogenic collapse. geological groups, under the leadership of arild andresen, have been active throughout the region between scoresby sund (71°29´n) and ardencaple fjord (75°30´n), and included participants from norway, denmark and the usa. the activities were funded by statoil and norwegian research foundations. in the early years activities were mainly in the central fjord zone (72°–75°n). in 2002 activities were concentrated on the arden caple fjord region, reached by twin otter with a landing at ny jonsbu. in 2003 a systematic collection was made of lamprophyres in the central fjord zone, and included visits to blomsterbugt, ella ø and strind berg land. in 2004 a party of 20 visited jame son land, milne land and kong oscar fjord. in 2003, 2004 and 2007 (and probably other years) arild andresen assisted a statoil group of geologists, by arranging a guided twin otter excursion. in 2007 activities were mainly on clavering ø and hold with hope. [dpc report archive.] 1996 stauning alper, nordaustgrønland (north-east greenland): frode guldal a norwegian expedition of nine persons led by frode guldal undertook a ski and climbing traverse of the stauning alper in april–may. a ski party was landed by twin otter in nordvestfjord close to the glacier løberen, and a climbing party on upper roslin glet scher. a total of 33 peaks were climbed, of which 16 were claimed to be first ascents. numerous passes were traversed, several for the first time. [dpc report archive.] 1996 mylius-erichsens mindeekspedition (mylius-erichsen memorial expedition): finn rasmussen a six-member expedition searched southern hov gaard ø and the islands and skerries off the front of nioghalvfjerdsfjord and lambert land for traces of the lost members of the danmark-ekspeditionen. no new relics were found. [dpc report archive.] 1996 the professor molchanov east greenland cruise the plancius foundation organised a cruise in late august with the cruise ship professor molchanov and 32 guests that called at foster bugt, myggbukta, antarctic havn, rypefjord, nordvestfjord and scores bysund. [dpc report archive.] 99 1996 scottish mountaineering club east greenland expedition: colwyn jones an eight-member expedition led by colwyn jones visited the stauning alper in july–august. they were landed by twin otter aircraft near majorpasset in the heart of the stauning alper. a number of peaks, in cluding dansketinde and hjørnespids, were climbed, and several first ascents were claimed including jaalspids (2100 m), susan’s peak (2238 m), aliertinde (2580 m) and annesketinde (2460 m). climbing was brought to a halt on 1 august by bad weather which forced a retreat to mestersvig that took eight arduous days (reid 1997). [dpc & rgs report archives.] 1996–98 geus studies of ice-sheet response to climate change in 1996 glaciological research was initiated by the geological survey of denmark and greenland (geus) and the danish polar center (dpc) on the floating glacier tongue filling nioghalvfjerdsfjord. the research was supported by the european com munity environment and climate programme. in 1997 quaternary field work was carried out around blåsø, a tidal lake at the margin of the floating glacier. in 1998 supplementary field work was undertaken on søndre mellemland and on île de france (now qa qer taq prins henrik; thomsen et al. 1997; bennike & weidick 1999). 1997–1998 geus geological mapping of the kong oscar fjord region the kong oscar fjord region from 72° to 75°n was mapped geologically as part of a regional mapping programme by the geological survey of denmark and greenland (geus) to produce 1:500 000 scale map sheets (escher 2001). survey geologists worked with an international group of guest geologists, under the overall leadership of niels henriksen (henriksen 1999). the main base was at mestersvig, with a secon dary base at krumme langsø. the mapping teams were supported by two helicopters, while twin otter operations were carried out under a charter agreement with the danish polar center that co-ordinated geus transport requirements with other expedition groups. this geological mapping project was notable for the first demonstration of large scale (hundreds of km) westward thrust displacement (higgins & leslie 2000; leslie & higgins 2008). 1997–present: zackenberg ecological research operations (zero) the ecological research station on the north side of tyrolerfjord, about 5 km north-east of the mountain zackenberg beside zackenbergelv was officially open ed in august 1997 after a two-year building phase. it initially comprised about 10 buildings, including laboratories, and had accommodation for 15 scientists. discussions on the possibility of establishing a permanent research facility in the north-east green land national park were initiated in 1986 (meltofte & thing 1996). a location in the daneborg– zacken berg region was considered appropriate as it lies in the transition zone between the lush and snow-rich southern parts of the high arctic and the more arid northern parts. the build ing phase was initiated in 1995, and included preparation of a 450 m runway suitable for twin-otter aircraft. the first zero annual report for 1995 was published in 1996 (meltofte & thing 1996). in 1994 a marine studies project was begun, based at the former weather station at daneborg. significant enlarge ments to the main facilities beside zacken berg elv were made in 2006–2007. the station is normally open from 1 june to 1 september, but in 2008 was extended from 13 march to 2 november. the total number of scientific visitors to the station were 81 in 2008, and 2700 overnight stays were re corded. the research station has a number of major re search programmes. the first observations for the ‘geobasis’ and ‘biobasis’ programmes were made as early as 1995, during the building phase. in 1998 a ‘klimabasis’ programme was added, and in 2002 the marine studies project started at daneborg in 1994 became the ‘marinbasis’ programme. in 2007 another major programme, ‘glaciobasis’, was initiated. in addition to the major research programmes, 10 or more large and small research projects are carried out each year (meltofte & rasch 2009). zackenberg ecological research operations (zero) was organised and operated by the danish polar center (dpc) up to 2008, but with the closure of dpc in early 2009 responsibility has been taken over by the national environmental research in stitute at aarhus university. zero issues annual re ports of activities, and in 2008 issued a thick sum mary volume recording 10 years of monitoring and research (meltofte et al. 2008). there are official place names for many of the mountains, rivers and valleys surrounding the re search station, but numerous unapproved names have 100 been introduced by visiting scientists for minor features, such as small lakes and areas of vegetation, and have been used in reports of their observations. 1998 british schools exploration society (bses) expedition to east greenland: pat cannings the british schools exploration society carried out a large-scale programme of adventure and exploration in the general mestersvig area. a total of 68 young adventurers were guided by 17 leaders under the overall leadership of pat cannings. [dpc & rgs report archives.] 1998–present: tangent expeditions tangent expeditions international (paul h. walker) began to organise climbing and ski expeditions to east greenland in 1989, and expanded activities to northern east greenland in 1998. expeditions that have given accounts of their activities in accessible reports are individually described under the year of activity below. those expeditions that were organised by tangent, or made extensive use of their logistic support, and that reached areas north of 69°n are briefly listed here. tangent has not deposited full reports of many of their expeditions with the danish polar center, apparently leaving this task to the initiative of individual leaders. many of the expeditions organised by tangent have been specifically aimed at ascents of unclimbed summits. in general, names given to summits where a substantial part of the ‘ascent’ was achieved by the use of twin otter aircraft are not included in this volume. 1998: three expeditions visited the petermann bjerg – shackleton bjerg region (73°n), and two the rigny bjerg region (69°03´n). 1999: expeditions were made to the lindbergh fjelde (69°07´n), to louise boyd land (73°30´n), and two to the rigny bjerg region (60°03´n). 2000: expeditions visited the lindbergh fjelde (69°07´n), dronning louise land (76°30´n), and two the rigny bjerg region (69°03´n). 2001: expeditions were made to the lindbergh fjelde (69°07´n), two to the rigny bjerg region (69°03´n, and two to the martin knudsen nunatakker (73°15´n). 2002: expeditions visited nils holgersen nuna takker (73°20´n) and louise boyd land (73°30´n), and two expeditions were made to both liverpool land (71°n) and knud rasmussen land (69°30´n); four expeditions touched on the gronau nunatakker (69°30´n). snow conditions in 2002 were reported as unusually poor. 2003: expeditions were made to liverpool land (71°n), the hvidbjørn nunatakker (73°38´n), the rigny bjerg region (69°03´n), and three to various parts of knud rasmussen land (69°30´n). 2004: expeditions visited liverpool land (71°n), knud rasmussen land (69°30´n), and two reached milne land (70°40´n). 2005: expeditions were made to liverpool land (71°n) and two to milne land (70°40´n). 2006: five expeditions visited parts of milne land (70°40´n), one liverpool land (71°n), two the re gion around sortebræ at 69°n, and one part of knud rasmussen land (69°30´n). 2007: two expeditions visited liverpool land (71°n), one making a n–s traverse. three expeditions visited dronning louise land (76°30´n), one went to a part of knud rasmussens land (69°30´n), and a large west lancashire scouts expedition climbed in renland (71°10´n; see report below). 2008: again three expeditions visited parts of milne land (70°40´n), two visited paul stern land (70°10´n), and an expedition led by georg czak made a ski traverse from near the watkins bjerge (69°n) to paul stern land (70°10´n; see report below). [information from ‘tangent expeditions/climb greenland’ website.] 1998 nunatak expedition: daniel caise a party led by daniel caise was landed by twin otter on 1 april at the mouth of knækdalen in inner kejser franz joseph fjord. they had as their main objectives the ascent of petermann bjerg and shackleton bjerg in april, but frustrated by poor snow conditions abandoned the knækdalen route and travelled south to hisinger gletscher, from where ascents were made of verena horn, hamlet bjerg and vähfreude. in kjerulf fjord the expedition observed a total of nine polar bears, including cubs. [dpc report archive.] 1998 ejnar mikkelsen fjeld expedition: h.c.f. sørensen an attempt on ejnar mikkelsen fjeld was made by a danish group led by h.c.f. sørensen in may–june, starting from scoresbysund and using skis and pulks to cross geikie plateau. 101 1998 swiss expedition to gunnbjørn fjeld: martin fischer a swiss expedition led by martin fischer made several ascents in the watkins bjerge in late april to early may. most activity was south of latitude 69°n, but a few days were spent farther north. 1998 rigny bjerg expedition: mark bailey a four-member mountaineering expedition led by mark bailey visited the rigny bjerg region in july. access was by twin otter aircraft. a total of 14 first ascents were claimed between 2000 m and 2600 m high, including an attempt on rigny bjerg (their mr. big) that turned back 183 m below the summit. [rgs report archive.] 1998 suess land kayak expedition: hugh simpson hugh simpson led a four-person expedition whose objective was to circumnavigate suess land. the intended portage of their kayaks along the north flank of hisinger gletscher proved too difficult, and the party returned to mestersvig. [dpc report archive.] 1998 scottish mountaineering club (smc) greenland expedition: colwyn jones colwyn jones led an eight-person smc expedition to the central stauning alper in may. the group flew in to constable pynt using a ski-equipped twin otter and, after refueling, continued to a base camp established on the upper part of sefström gletscher. several first ascents were made, and on 15 may the party began an arduous six day journey through deep snow to reach mestersvig on 21 may. [dpc & rgs report archives.] 1998 vertebrate palaeolontological expedition to jameson land: farish jenkins farish a. jenkins jr. led a 10-person expedition to north-east jameson land, continuing his earlier in vestigations (1988–89 & 1991–92) of vertebrate fossils in triassic sediments. a further expedition was made in 2001. [dpc report archive.] 1998–99 øfjord expedition: grundtvigskirken grundtvigskirken, a spectactular mountain on the north-west side of øfjord dominated by a central granite tower 1997 m high, was the objective in 1998 of a four-person climbing group (three norwegians and one swedish member) led by bengt nilfors. the approach was made using kayaks. a further attempt in 1999 by the same group (but with a different swedish climber) was successful. the climb by the south ridge took 2½ days, and one of the particants commented that ‘it was the best alpine rock climb he had done’ (anonymous 2000 p. 241). a small cairn on the summit recorded an earlier ascent by an easier route, probably the south-west face. in their report the mountain is named tsavagattaq, a greenlandic name for the tip of a harpoon. in 2010 the mountain was climbed by a ‘national geographic’-supported climbing party (hans ambühl, personal communication 2010). 1998–present: nanu travel aps the icelandic travel company nonni travel founded a branch at scoresbysund / illoqqoortoormiut (ittoq qor toormiit) in 1998, later changing its name to nanu travel aps. this company has greatly assisted the promotion of tourism in this part of east greenland, and in particular has brought the visits of cruise ships to the town into a comfortable routine to the benefit of both the visitors and the resident population. cruise ships that visit east greenland most frequently call at scoresbysund / illoqqoortoormiut (ittoqqortoormiit) on their passage from longyear byen (svalbard) to keflavik (iceland), or vice-versa. the ships used by the shipping companies have a degree of ice-strengthening to ensure safe passage through the ice-belt, and carry scientific experts who act as guides. in recent years the shipping lines most active have included: oceanwide, peregrine ship ping, aurora expeditions, polar star expeditions, albatros, phoenix and quark. depending on ice conditions, the ships may call at historically interesting sites in the north-east greenland national park, or localities where musk oxen and other wild life can be viewed safely. nanu travel aps at scoresbysund / illoqqortoormiut (ittoqqortoormiit) have recorded up to 17 visits annually by cruise ships in recent years, carrying an annual total of 800–1000 passengers. [nanu travel, personal communication 2008.] 102 1999 swedish øfjord expedition a party of seven swedish climbers visited the øfjord region of the inner scoresby sund region, and climbed two of the summits of a mountain group about 4 km south-west of grundtvigskirken. in their report this mountain group is erroneously assumed to be grundtvigskirken, although it has no resemblance to the church grundtvigskirken in copenhagen. a party of four climbed the middle summit by the south-east pillar, and the south face of the southern spire was climbed by two members in 25 pitches (ano nymous 2000). 1999 ‘arcturus’ clavering ø expedition: simon fraser an expedition organised by the travel company arcturus, and led by simon fraser, visited clavering ø in july–august. natural history observations of birds and plants and visits to inuit archaeological sites were made. 1999 tangent rigny bjerg expedition a four-person expedition led by nigel edwards, and organised by tangent expeditions, visited the rigny bjerg region in late may, and claimed 10 first ascents (gregson 2000a). 1999 greenland rigny bjerg mountain eering expedition a three-man party explored part of the rigny bjerg region from 3 to 25 july, and ascended three peaks. a base camp was established by twin otter at 69°18´n, in co-operation with tangent expeditions. the weath er was clear, sunny and calm, with temperatures around minus 20°c. in addition to their three ascents, the party also carried out six two-day exploratory ski tours (mitchell 2000). [rgs report archive.] 1999 tangent expedition to the lindbjergh fjelde paul walker of tangent expeditions led a nine-per son expedition to the lindbergh fjelde region (69°n 31°w) from late july to early august. access was by twin otter aircraft. the rock was reported to be of poor quality, but the weather was perfect and 20 ascents between 2600 and 3200 m were made in 16 days (gregson 2000b). 1999 scoresbysund ecological studies: hans-ulrich peter ornithological and associated botanical and biological studies were carried out by a small group in july– august around scoresbysund; the group was led by hans-ulrich peter. 1999 cambridge north-east greenland expdition: mathew tinsley a group of five led by matthew tinsley visited louise boyd land from 2 july to 26 august. basecamp was established at 73°30´n 28°00´w, and the climbs around the base were mostly ski ascents with little technical difficulty. north of base a secluded high glacier gave access to more challenging climbs on excellent granite. on their way south to their pick-up point, petermann bjerg was climbed by two new routes, thought to be the sixth and seventh ascents. their descent to the pick-up point at the head of kjerulf fjord involved an awkward descent of hisinger gletscher and a 5 km hike along bocksriet dalen (bostock 2000). [rgs & bmc report archives.] 1999 young sund walrus studies: erik w. born erik w. born (greenland fisheries research insti tute) carried out studies of the energy requirements of walrus in young sund, east greenland. sandøen in young sund is one of the regular haul-out localities for walrus (born et al. 1997). 1999–2000 maria ø expedition: john thorogood john thorogood led three-person expeditions to north-east greenland in 1999 and 2000, using an inflatable boat for transport in the fjords. 1999 – boat trips were made in the fjord system east and west of ella ø and maria ø in july–august. ascents were made of angelin bjerg and rødebjerg on ymer ø. boat trips were also made to the heads of rhedin fjord, röhss fjord and dickson fjord. an ascent was made of langenthaler gletscher, to the col overlooking concordiaplads. 2000 – boat visits were made to the inner parts of alpefjord and forsblad fjord in august. [dpc report archive.] 103 2000 caledonian eclogite studies: jane gilotti jane a. gilotti continued her studies in the northern east greenland eclogite province in july. her earlier studies were carried out as part of the ggu geological mapping of the dove bugt region in 1988–90, her own expedition to danmarkshavn in 1992, and as a member of the ggu/geus geological mapping group in the lambert land region in 1993–1995. 2000 geologfjord expedition: iain smart a journey by inflatable boat was made from mesters vig to geologfjord in august–september, led by iain smart, in association with nanu travel aps. 2000 rigny bjerg – watkins bjerge expedition an eight-member british party led by brian needle man visited the rigny bjerg region and made a num ber of first ascents. two members of the group skied west to the watkins bjerge, and made a successful ascent of gunnbjørn fjeld (gregson 2001a). 2000 scottish suess land expedition: douglas anderson boat tours in the fjords around suess land, lyell land and frænkel land were made in july–august, by a party led by douglas anderson. 2000 hvalrosundersøgelser i nordøst grønland (walrus studies in north-east greenland): erik w. born erik w. born returned to lille snenæs (76°53´n) on the south coast of germania land to continue his studies of walrus at the same location in 1989–90. this was part of a project ‘changes in arctic marine production’ supported by grønlands naturinstitut (greenland nature institute) and grønlands miljø undersøgelser (greenland environmental investi gations). e.w. born and m. acquarone were flown to hvalrosodden by twin otter, from where they pro ceeded to lille snenæs; they took blood samples from seven walruses and attached radio senders to six walruses during the summer. on their return to hvalrossodden in late august the party observed several walruses hauled out on the hvalrosodden peninsula, site of the massacre of 11 walruses during the 1906–08 danmark-ekspeditionen (born & acqua rone 2001). 2000 graae-rasmusen expedition walking tours and visits to danish hunting huts in hochstetter forland were made in august by a twoman party, jesper graae and hans rasmusen. [dpc report archive.] 2000 lancaster university hiking club expedition to dronning louise land ski touring and mountaineering were undertaken in the nunatak region of dronning louise land in may by a nine-person group from lancaster university hiking club. there were delayed for about a week at constable pynt due to rescue operations for an expedition that had lost a man down a crevasse. the expedition was eventually flown in to dronning louise land by twin otter. twenty-two summits were climbed, but many of them were small nunataks only a few hundred metres above the surrounding ice cap surface. the rock was mainly of poor quality and ascents presented little difficulty. 2000 british dronning louise land expedition scott umpleby led a climbing group to dronning louise land, but like the lancaster university expedition (see above) they were delayed at constable pynt. when they eventually arrived by twin otter at dronning louise land, they set up their base camp at 1870 m, and climbed 34 summits in 10 days (gregson 2001b). 2000 expedition to lindbergh fjelde: paul walker paul walker (tangent expeditions) led a 10-person group to the lindbergh fjelde, where a base camp was established at 2120 m by twin otter aircraft. about 16 summits were climbed, ranging in altitude from 2260 to 3150 m (keaton 2001). 2000 expedition sirius teams of the sirius sledge patrol undertook a fourmonth sledge journey from thule (qaanaaq) in north-west greenland across north greenland and down the coast of northern east greenland to dane borg (expedition sirius 2000). this particular journey by sirius deserves special note only because the members of the patrol included crown prince frede 104 rik of denmark, and the activities were given wide press and television coverage. 2000 late quaternary history of jameson land: lena andrielsson lena andrielsson led a group of four to the ugleelv area of jameson land in july–august, investigating late quaternary deposits. earlier studies in the same area were made during the ‘1990–1992 ponam (polar north atlantic margins) project’. 2000–2004 sports-ekspedisjon i nord østgrønland (sports expedition to north-east greenland): hans lapstun the norwegian hans lapstun undertook a series of summer sports expeditions to northern east green land. [dpc report archive.] 2000 – walking tours were undertaken in the mestersvig area with a small group in july and august. 2001 – boat tours were made with a small group in the fjords north of mestersvig, visiting ella ø, blom ster bugten and strindberg land. 2002 – a zodiac rubber boat was used for transport through the central fjord region, calling at kap peterséns, blomsterbugten, strindberg land, ella ø, sorte hjørne and nyhavn. lapstun’s report contains information on the condition of the huts he visited, and notes minor repairs he carried out. 2004 – a three-man group led by hans lapstun under took a three week tour on foot between nat horst fjord and mestersvig in july–august. 2001 smog in greenland: martin knudsen nunatakker: mark lampard an eight-person slough mountaineering group (smog) party led by mark lampard visited the mar tin knud sen nunatakker (73°15´n) in june, undertaking skitouring and climbing. poor weather meant they were landed by twin otter 40 km short of their destination, but this was reached with a three-day sledge trip and 30 summits 2100–2700 m high were climbed. [bmc & rgs report archives.] 2001 ‘quest’ historisk expedition (‘quest’ historical expedition): jan brun a norwegian tourist expedition with 43 participants led by jan brun sailed with the fogo isle to east greenland in august. landings were made at the former norwegian and danish hunting stations at germaniahavn, kap herschell, revet, krogness, mygg bukta and ella ø. [dpc report archive.] 2001 lanchester greenland expedition: jonathan white jonathan white led a six-member party on a climbing expedition to the lindbergh fjelde. access was by twin otter aircraft, and the landing site on 22 june was at 69°07´n 31°02´w. all six members climbed their first peak, after which the party split into groups of two or three. a total of 28 summits were climbed, ranging from 2270 m to 2935 m in altitude, of which 25 were thought to be first ascents (white 2002). on one summit a survey pin was found drilled into the rock, but this was not, as surmised a relic of martin lindsay’s surveying in 1934, but a fixed point established by the danish geodætisk institut (geodetic institute) in the mid-1980s (willy weng, personal communication 2004). [bmc & rgs report archives.] 2001 vertebrate palaeolontological expedition to jameson land: farish a. jenkins farish a. jenkins jr. led a four-person expedition to north-east jameson land in june–july, to continue his earlier investigations (1988–89, 1991–92, 1998) of vertebrate fossils in triassic sediments. [dpc report archive.] 2001 bioteknologisk institut østgrønland expedition (danish technological institute, east greenland expedition): peter stougaard a three-person group from this institute based in hørsholm, denmark, made investigations of the hot springs on liverpool land and the northern blosse ville kyst. nørrefjord, rømer fjord and knighton bugt were visited. [dpc report archive.] 2001 scottish mountaineering club expedition: colwyn jones colwyn jones led a six-man party to the central stauning alper in july–august. from mestersvig the party was lifted by helicopter to the upper part of cantabræ. the weather was generally fine and stable, and climbs were made of sussex (2330 m) and first ascents of pap of cumbrae (1885 m), tandlaegetinde 105 (2350 m), keswicktinde (2380 m) and mears fjeld (2100 m) (read 2002). the return to mestersvig was also by helicopter. [bmc & rgs report archives.] 2002 øst-grønland under seil (east greenland under sail) the 33-foot sailing boat lodyn with a three-person norwegian crew sailed from bergen via iceland to reach scoresbysund/illoqqortoormiut on 2 august. they travelled into the fjord making landings at sydkap, the bjørnøer and bregnepynt before returning to scoresbysund. on 12 august they left green land on their return voyage. [dpc report archive.] 2002 nils holgersen nunatakker expedition: paul walker a party of six climbers led by paul walker (tangent expeditions) left iceland by twin otter on 17 june, and after refueling at constable pynt landed in the nils holgersen nunatakker (73°20´n). the weather was perfect, with unexpectedly high temperatures. about 16 peaks ranging from 2061 to 2543 m high were climbed, some involving 7 km long ski journeys on the glaciers. the party departed by twin otter on 5 july (keaton 2003). 2002 gronau nunatakker british-american expedition a group of six climbers visited the gronau nuna takker in july, establishing a base camp at 69°28´n by twin otter aircraft. they skied and climbed in two groups, and made a series of ascents of peaks up to 2900 m high (burch 2003). 2002 with the arnak from mestersvig to strindberg land the arnak is a 23-foot motor-cutter built in 1963, and from 1967–1997 used by sirius to lay out depots in the fjord region of east greenland. in 1997 it was purchased by two former sirius members, and is stationed at mestersvig in the winter. a 2002 voyage in the southern part of the north-east greenland national park from mestersvig to strindberg land and return is described by christensen (2003). 2002 exploration of north-west watkins mountains, east greenland a two-person group (al read and john hulse) were dropped off on the ice cap in the north-west watkins bjerge by twin otter, and in june and july explored and climbed four easy summits north of 69°n. the summits appear in their report as summit 1 to summit 4. 2002 cambridge greenland glaciology expedition: chris lockyear a five-person expedition led by chris lockyear visited northern louise boyd land in july and august. the expedition was landed at their study area by twin otter. glaciological and geological studies were carried out, and six climbs were made of summits up to 2340 m high (lockyear 2003). following completion of the scientific work, a long 17-day ski traverse was made west of louise boyd land, along victor madsen gletscher, west of the heads of nordenskiöld gletscher and hisinger gletscher, and via langen thaler gletscher to dickson fjord (72°50´n) where they were picked up by rubber boat. [dpc report archive; bmc report archive.] 2002 belgian expedition aboard the kittiwake this expedition reached greenland on 17 august, an choring at ella ø. the ship then sailed into the inner ends of kejser franz joseph fjord and kjerulf fjord, returning via antarctic sund and kong oscar fjord to forsblad fjord where the ship anchored at the western inner end. a short visit was made to scoresbysund before leaving for europe on 3 september. [dpc report archive.] 2002 liverpool land and knud rasmussens land: tim mosedal tim mosedal with a small group visited southern liverpool land and parts of knud rasmussen land. a total of 15 peaks were climbed. 2002 loughborough grammar school greenland expedition an expedition comprising 13 pupils and two teachers from loughborough grammar school (uk), and two 106 leaders from tangent expeditions, visited liverpool land in july. six first ascents were claimed of summits west of bjerring pedersen gletscher. [rgs report archive.] 2002 shackleton bjerg expedition: john thorogood this four-person expedition was led by john thoro good, and travelled by boat from mestersvig to the head of dickson fjord. they reached the ice cap by ascending langenthaler gletscher on gletscher land and climbed shackleton bjerg and several nearby summits including verena horn and guld tinderne. [dpc report archive.] 2002 ‘explorers corner’ north-east greenland national park sea-kayaking trip: olav malver olaf malver led a 12-member group on a kayaking tour north of mestersvig in august. the group was flown back to mestersvig from strindberg land. [dpc report archive.] 2002 international geological expedition to jameson land: jennifer mcelwain an international group of geologists visited ranun kel dal and astartekløft in jameson land in july– august. this field work was part of a project ‘fossil floral dynamics across the triassic-jurassic boundary of east greenland’ funded by the national geograp hic society of the usa. the expedition was extremely successful. [dpc report archive.] 2002 french literary expedition a five-person expedition sailed through the central fjord zone using three kayaks and a rubber boat for transport. the objectives were to make a film to celebrate the life of jørn riel, a noted danish author who took part in lauge koch’s geological expeditions in east greenland in the early 1950s (rohan et al. 2003). jørn riel is noted in particular for his ‘skrøner’, a series of ‘tall stories’ or fables loosely based on his greenland experiences. 2002–03 jónas g. allanson visit to scoresbysund jónas g. allanson stayed at scoresbysund / illoqoor toor miut (ittoqqortoormiit) from september 2002 un til the summer of 2003 as part of a research project on the use of marine resources by an isolated community. he took part in local hunting journeys. [dpc re port archive.] 2002–2007 ‘arcturus’ expeditions to north-east greenland a series of expeditions led by robert burton and kath leen cartwright, and organised by the travel com pany arcturus, visited various parts of northern east greenland. [dpc report archive.] 2002 – robert burton and kathleen cartwright led a 12-person expedition to wollaston foreland in july–august. the party were landed by twin otter at slettedalen, and made observations on the fauna, flora and archaeological sites. in august robert burton led a three-person expedition to the region around blyklippen near mestersvig in august, mainly bird-watching and scrambling. washburn’s hus west of mestersvig was used as a base. 2003 – an eight-person expedition with archaeological objectives made observations in the inner fjords of the scoresby sund region in august. amongst other places, hekla havn and c. hofmann halvø were visited. another 12-person expedition led by kathleen cartwright and robert burton, visited the clavering ø region in july. they used a twin otter aircraft to reach revet, and a zodiac rubber boat for local transport as far as eskimonæs. archae ological and botanical observations were carried out. 2005 – a 15-person expedition, led by kathleen cartwright and robert burton, was transported to southern geographical society ø in july by twin otter. one group surveyed inuit ruin sites on the shore of vega sund. 2006 – this expedition led by kathleen cart wright visited the southern part of dove bugt in july and august, landing by twin otter at the airstrip adjacent to the sirius hut in ravnedal, rechnitzer land. botanical observations and registration of sparse inuit ruins were made in the coastal areas of rechnitzer land and on the north coast of ad. s. jensen land to the south-east, reached using inflatable rubber boats. 2007 – an eight-person expedition led by kath 107 leen cartwright explored the coastal region of eastern lyell land in july and early august, landing by twin otter on a rough landing strip between kap lager berg and kirschdalen. archaeological sites were visited and surveyed, including some on åkerblom ø and at kap harry on ella ø, reached by zodiac rubber boat. 2003 expedition to knud rasmussen land, east greenland a six-person expedition visited the northern watkins bjerge in the region formerly known as knud ras mussen land in july and august, landing by twin otter aircraft at 69°08´n. a total of 20 first ascents were claimed. [rgs report archive.] 2003 bses liverpool land expedition the british schools exploration society (bses) took a large party of young people to liverpool land in july and august. the overall leader was john muston, and 15 deputy leaders were in charge of the 60 young explorers divided into five groups. access was via constable pynt. [rgs report archive.] 2003 ecopolaris (grea) this 2003 ecopolaris expedition, part of the arctic activities of groupe de recherches en écologie arc tique (grea – arctic ecology research group), visited north greenland and northern east green land, and amongst other activities ringed 270 ivory gulls. on henrik kröyer holme inuit ruins were inspected, and in dove bugt new walrus haul-out locations were recorded. [dpc report archive.] 2003 liverpool land geological studies: ebbe hartz ebbe hartz visited liverpool land in august, and collected samples for isotopic age determinations around storefjord and mariager fjord. samples were also collected for ebbe hartz at different altitudes on dansketinden by stephen reid’s scottish mountain eering club expedition (see entry below), in exchange for helicopter transport. [dpc report archive.] 2003 scottish mountaineering club east greenland expedition: stephen reid a four-man expedition led by stephen reid visited the stauning alper, starting with a helicopter lift to majorpasset (col major) at the foot of dansketinden, the central focus of the summer’s climbing. despite periods of poor weather, new spectacular routes were made on the south and south-west ridges of danske tinden. [bmc, dpc & rgs report archives.] 2003 ‘tunu-i’ expedition: fish fauna north-east greenland fjords a ship-based expedition aboard the jan mayen visited the fjord region between 74° and 77°n in october, to study the fish fauna of the fjords. this expedition was planned following a brief visit to dove bugt and godthåb golf in october 2002, and further expeditions are planned. [dpc report archive.] 2003 ‘midnight sun 03’ expedition to rigny bjerg: martin bohl an eight-person expedition to the rigny bjerg area led by martin bohl visited the rigny bjerg area in july, and claimed to have climbed 38 summits. one of these was the 1971 m high peak wrongly identified as ‘rigny bjerg’ on ams (american map service) maps, but on 19 july 2003 martin bohl and mike palmer climbed the real rigny bjerg, the highest summit in the area. [bmc & rgs report archives.] 2003 nordvestfjord kayak expedition a small group of kayak enthusiasts visited the inner parts of nordvestfjord in july–august, reaching as far west as eskimovig. [dpc report archive.] 2003 nord-østgrønlandsk kayakekspedition (north-east greenland kayak expedition) this four-person expedition visited the fjord region of north-east greenland in july and august. starting from daneborg, they travelled around clavering ø to the head of loch fyne, made a long portage through stordal to reach the head of moskusoksefjord, and continued via ella ø to mestersvig. [dpc report archive.] 108 2003 geological expedition to jøkelbugten, north-east greenland: jane gilotti jane a. gilotti led a four-person geological expedition to the jøkelbugten region to continue her studies of eclogites. investigations were mainly carried out in sanddal reached by twin otter aircraft, and at rabbit ears island in jøkelbugten reached by helicopter. this was a continuation of her earlier eclogite studies with ggu/geus mapping expeditions (1988–90, 1993– 1995) and her own expeditions in 1992 and 2000. [dpc report archive.] 2003–present: geocenter danmark – east greenland activities geocenter copenhagen was established in 2003 as a formalised cooperation between the geological sur vey of denmark and greenland (geus), the danish lithosphere centre (dlc), the geological museum and the geological and geographical institutes of the university of copenhagen. dlc closed down when its funding expired in 2004. in 2008 with the inclusion of the department of earth sciences of the university of aarhus the name was changed to geo center denmark. annual summaries of field activities planned in greenland by the geocenter partners were issued from 2004 onwards. many geographers and some geologists have been attached to the ecological projects operated by the zackenberg ecological research operations – zero (see above, ‘1997– pre sent zackenberg ecological research opera tions’) and are included in the annual sum maries. extracts of work planned in northern east green land (excluding zero projects) are given below. 2004: the geological museum continued studies of cambro-ordovician sediments on ella ø. in southeast jameson land studies were made of the kap stewart formation, with special reference to the mass extinction that took place at the trias/jurassic boun dary. 2005: ella ø was again the subject of studies, but of the eleonore bay supergroup and tillite group, with collection of material for analyses of stable car bon isotopes. geus carried out studies in the mesters vig region in july and august, a follow-up of hyper spectral anomalies discovered during airborne surveys in 2000. 2006: studies of the neoproterozoic eleonore bay supergroup and tillite group were extended to andrée land, with collection of material for analyses of stable carbon isotopes to be compared with results from samples taken in 2005 on ella ø. a small group from geus undertook sedimentological and stratigraphical studies of late carboniferous sediments in two coast profiles in eastern kronprins christian land. 2007: neoproterozoic–palaeozoic geological studies were continued on ella ø, with an extension of the study region to krumme langsø. lower cretaceous rocks on wollaston forland were investigated. 2008: northern east greenland was the focus of a variety of studies in 2008. monitoring of the inland ice (dancea project; danish cooperation for environment in the arctic) involved establishment of automatic mass-balance stations on the inland ice margin in kronprins christian land, on violin gletscher in nathorst land and in a.p. olsen land near zackenberg. another mass-balance station was established near malmbjerg by quadra mining, in connection with the planned mining project. 2009: the automatic mass-balance stations established in 2008 on the inland ice margin in kronprins christian land and on violingletscher, and on local ice caps in a.p. olsen land, were inspected and necessary maintenance carried out. the station erected by quadra mining on schuchert gletscher was also visited. geus undertook a major project, led by jørgen bojesen-koefoed, involving seven field teams working between jameson land in the south and germania land in the north, aimed at the sedimentology and oil geology of rock sequences related to the offshore sedimentary basins; a special group undertook drill coring of jurassic and cretaceous sequences in wollaston foreland and hold with hope. ten seismological stations were established along a 250 km e–w cross-section of the caledonian orogen at about 73°30´n (aimed at the registration of distant natural earthquakes) by a small group from the department of earth sciences at aarhus university; samples for fission-track analysis were also collected. trine dahljensen co-ordinated activities by geus and the institute of geography and geology at copenhagen university, that undertook the establishment of 22 seismometers along a 610 km profile at 70°n; mea sure ments over a period of two years will contribute to new detailed models of the earth’s crust and upper mantle. [geus archive.] 2004 sailboat jonathan visit to ne greenland the sailing boat jonathan with a crew of two sailed 109 from longyearbyen (svalbard) to the northern east greenland fjord region, where they visited vega sund, geologfjord, ella ø and mestersvig. [dpc report archive.] 2004 liverpool land ski-mountaineering a four-person group visited liverpool land in april–may. from scoresbysund/illoqqortoormiut they reached their starting point by dog sledge, and spent more than three weeks climbing and skiing around emmanuel gletscher, kolding gletscher and åge nielsen gletscher before returning to scores bysund (thomson 2005). [climb magazine news letter, november 2005.] 2004 west lancashire county scouts mountaineering group east greenland expedition a large group of scouts from west lancashire visited milne land, reached by twin otter aircraft, and climbed numerous peaks on both sides of korridoren, the large glacier that cuts across the island from east to west. this well-organised expedition led by dick griffiths made a number of impressive ascents (griffiths 2004). a selection of unapproved names given to peaks is included in this volume. [dpc & rgs report archives.] 2004 chicago field museum expedition to kap stewart a three-person expedition from chicago field mu seum visited kap stewart in july–august to collect triassic–jurassic fossils. [dpc report archive.] 2004 british expedition to knud rasmussen land a four-person british expedition flew by twin otter into the region south of scoresby sund formerly known as knud rasmussen land, and established a base camp at 69°10´n. a total of nine first ascents were made, up to 2884 m high (windsor 2005). [climb magazine newsletter, november 2005.] 2004 rando aux alpes de stauning (french climbing expedition to the stauning alper) a seven-member expedition led by yves dupont visited borgbjerg gletscher in the south-west stauning alper in april. they made slow progress in deep and sticky snow, abandoned attempts at climbing and returned to constable pynt. [dpc report archive.] 2004 ecopolaris (grea) tara 5 expedition to ne greenland this was a more ambitious and wide-ranging tour than the usual land-based groupe de recherches en écologie arctique expeditions (grea – arctic eco logy research group). the tara 5 is a 36-metre iceclass schooner built in 1989, with a crew of five and space for about 15 scientists and their equipment. the expedition carried out investigations between the blosseville kyst (69°n) and danmarkshavn (76°46´n) between 8 july and early september, before sailing eastwards to jan mayen and south to iceland. akureyri was reached on 7 september. [dpc report archive.] 2005 greenpeace ship expedition: project thin ice: martina krüger martina krüger led an expedition on the ship arctic sunrise that visited the inner part of nordvestfjord. a survey was made of parts of daugaard-jensen glet scher and f. graae gletscher, and an iceberg survey was conducted in scoresby sund. a short visit was also made to the zackenberg research station. [dpc report archive.] 2005 cruise of grigoriy mikheev: dennis schmitt discovers new island in liverpool land in 2005 dennis schmitt was aboard a cruise with the grigoriy mikheev and in early september reported sailing around an island in northern liverpool land that was not marked on the map and was unknown to the residents in scoresbysund. this island has been unofficially named uunartoq qeqertaq / warming is land. the new island was widely reported in the international press as dramatic evidence of climate warming, but is in fact the result of slow melting of a small ice cap over a period of 10–15 years. [dpc report archive.] 2005 east greenland ship visits in addition to the 17 visits by regular cruise ships, 110 nuna travel aps recorded visits by several small ships to scoresbysund / illoqqortoormiut (ittoqqor toor miit) in 2005. these included the coelan, vagabond and vamos. [nanu travel aps personal communication 2008.] 2005 ‘tunu-ii’ expedition: fish fauna of north-east greenland fjords this continuation of the 2003 ‘tunu-i’ investigations had been planned for the region between bredefjord (75°33´n) and carlsberg fjord (71°26´n), but the heavy pack ice in late september prevented access, and activities were diverted to scoresby sund. trawling was carried out by the ship jan mayen at nine locations, and hydrographic stations were established in føhnfjord and at kap stephenson. [dpc report archive.] 2005 east milne land expedition: barry roberts barry roberts led a group to milne land for tangent expeditions. the snout of charcot gletscher was reached after a seven hour journey across the sea ice from constable pynt. about 20 ascents were made, mainly by ski. the expedition flew out by twin otter aircraft. some of the climbs made were repeated by the ‘2006 milne land expedition’ led by phil poole (editors comments in poole 2007). 2005 comer scientific studies of glacial features and relative sea-level changes the past 12 000 years an expedition aboard the r/v turmoil, equipped with a helicopter, visited the scoresby sund region in august. two field parties were set out in the schuchert dal area, and visits were made by helicopter to milne land and the stauning alper. [dpc report archive.] 2005–06 trekking expeditions to milne land: jim gregson jim gregson led trekking expeditions to milne land for tangent expeditions in both 2005 and 2006. in 2005 a group visited the arabertoppen area of southeast milne land. in 2006 a group visited south-west milne land where a number of easy ascents were made at about 70°25´n 27°49´w. access required an uncomfortable trip of 200 km in an open boat through heavy pack ice (gregson 2007a). 2005–08 malmbjerg new exploration phase quadra mining ltd. initiated a new phase of exploration and drilling at the molybdenum prospect at malmbjerg, following a dramatic increase in metal prices. this major activity undertook extensive new drill-coring. the project was put ‘on hold’ in 2008 when prices collapsed at the beginning of the financial crisis. 2005–present: british north-east greenland project in a continuation of the earlier project of the same name, small expeditions visited areas for walking tours and natural history observations, using inflatable boats for local transport. 2005 – visit to krumme langsø and the men an der øer. 2006 – alpefjord was visited. 2007 – a small group led by michael j. lea flew to krumme langsø by twin otter, and visited sur rounding areas using a rubber boat for local transport. 2008 – vega sund. a seven-person expedition led by michael j. lea visited the vega sund region in july and august, flying in and out with twin otter and using rubber boats for local transport. numerous musk oxen were seen on geographical society ø and single, non-aggressive polar bears were encountered. [dpc report archive.] 2006 scoresby sund late glacial ice advances: meredith a. kelly a seven-person group from several usa universities made investigations of glacial advances in the scores by sund region in western jameson land and on eastern milne land in august 2006. camps were set out and moved by the helicopter based at con stable pynt. [dpc report archive.] 2006 oxford university greenland expedition to gronau nunatakker: hauke engel a three-person expedition flew by twin otter from iceland direct to their target area at around 69°30´n in the gronau nunatakker on 6 august. they claimed 111 12 first ascents. they flew back to iceland on 29 august, on a shared charter with the anglo-scottish expedition (engel 2007). [rgs report archive; climb magazine newsletter, january 2008; bmc report archive.] 2006 milne land circumnavigation by kayak a three-person group flew into constable pynt and chartered a boat to carry them and their kayaks to south-east milne land. from here they travelled along føhn fjord and on south-west milne land climbed a rock route to hermelintop (called ‘hergen litop’ in their report). after completing their tour around milne land they paddled their kayaks back to scoresbysund/illoqortoormiut (ittoqortoor miit) (san ders 2007). [climb magazine newsletter, january 2008.] 2006 tangent expeditions visits to knud rasmussen land, sortebræ mountains two groups from tangent expeditions were active in the sortebræ region in may 2006. a rosie goolden group arrived in early may and spent 20 days in the region, making a number of first ascents at the mar gin of the ice cap. the group was lifted out on 27 may by the twin otter that brought in a six-person party led by jim gregson. this group established a base camp at 69°05´n, and made six first ascents up to 2405 m high (gregson 2007b). 2006 ‘brathay’ expedition to knud rasmussen land: paul williams an eight-person expedition from brathay explora tion group, led by paul williams, made a visit to a group of nunataks in the western part of geikie plateau from 17 july to 7 august. four peaks were climbed ranging from 1950 m to 2350 m high. ten rock samples with lichen were collected for the university of copenhagen (griffin 2007). [climb magazine newsletter, january 2008.] 2006 milne land expedition: phil poole phil poole led a three-person expedition to milne land from 8 to 16 may, reached by skidoo from constable pynt. a base camp was established about 10 km up char cot gletscher, and ski ascents were made of seven peaks from 1254 m to 1770 m high. the expedition was lifted back to constable pynt by helicopter because of the melt (poole 2007). 2006 bses liverpool land expedition the british schools exploration society (bses) again organised a trip to liverpool land. the 14 leaders and 54 young explorers were landed at constable pynt, and used local boats to reach a base camp established near the head of hurry inlet at kalkdal. the five groups of young explorers ranged northwards as far as carlsberg fjord and south to sødal carrying out a variety of scientific projects. [dpc report archive.] 2006 glasgow academy expedition to milne land: neal gwynne a 16-member expedition of four leaders and 12 pupils from glasgow academy (scotland) visited milne land in july. the group, led by neal gwynne, was landed by twin otter on eastern milne land, and walked south and west to reach the upper reaches of charcot gletscher. a series of peaks up to 1800 m high were climbed. two first ascents were claimed, while many other summits were reported to have previously been climbed by a ski-tour expedition. [dpc & rgs report archives.] 2006 anglo-scottish expedition to the wager nunatakker and watkins bjerge this four-person expedition was dropped off in the wager nunatakker at 69°39´n 27°44´w by the twin otter taking out the ‘brathay’ expedition. over the next two weeks the group undertook ski-mountain eering and alpine mountaineering. after a long icecap crossing, they were picked up in the northern watkins bjerge, together with the three members of the oxford university expedition. a total of 16 sum mits were climbed, mostly first ascents, but few of the peaks were more than a few hundred metres above the ice-cap surface; while providing spectacular views in a very isolated region none of the climbs were apparently of great difficulty. [bmc, dpc & rgs report archives.] 2007 smc east greenland expedition: colwyn jones an eight-member expedition from the scottish mountaineering club (smc) led by colwyn jones 112 visited the stauning alper in april and may. after landing by twin otter on the upper part of stor gletscher, a series of first ascents were made on both the west and east sides of the glacier. the expedition then moved northwards, climbing several summits around the upper gullygletscher, then crossed major passet (col major) and descended bersærker bræ to eventually reach mestersvig. many of their peaks were given danicised names, ‘spids’, ‘tinde’ and ‘bjerge’, although the singular form ‘bjerg’ would have been more accurate. [rgs report archive.] 2007 west lancashire county scouts mountaineering group east greenland expedition a large group of scouts from west lancashire visited renland, established a base camp on edward bailey gletscher, and made numerous climbs of peaks and high points on the ice caps to the north, south and west (griffiths 2007). access was by twin otter to an established rough airstrip on eastern milne land, from which speedboats were hired to ferry the group to the coast of renland. some helicopter transport was also necessary. like the earlier 2004 expedition to milne land, this was a well-organised and successful expedition that achieved all its objectives. unap proved names were given to 34 summits climbed; only a selection of names are included in this volume. [bmc, dpc & rgs report archives.] 2007 army mountaineering association: ‘boreal zenith’ expedition to andrée land: sam marshall mountaineering was carried out in july by a party led by sam marshall from a base camp established in central grejsdalen, reached by twin otter aircraft. the members of the party made numerous climbs of summits to the north and south of the valley, most claimed as first ascents. the expedition made use of geodætisk institut 1:250 000 scale topographic maps compiled in 1932 being unaware of the existence of modern topographic maps. of the 25 summits climbed and named by the boreal zenith expedition, only selected names are included in this volume. [bmc & dpc report archives.] 2007 british dronning louise land expedition this three-member expedition led by gavin booth visited dronning louise land in may–june. ten nunataks were climbed, of which eight were thought to be first ascents. a twin otter aircraft was used for transport. [rgs report archive.] 2007 japanese milne land expedition: yasushi yamanoi the japanese mountaineer yasushi yamanoi made an aerial reconnaissance of milne land looking for a suitable mountain wall to climb. he returned in july with his wife taeko and two others, one a tv producer with the japan broadcasting company. the party took a helicopter from constable pynt to ittoqqor toor miit from where they hired a boat to take them and their climbing and film equipment on an eight hour journey to east milne land. another helicopter lift on 27 july took them to the foot of a 1250 m vertical wall, that they named orca, at the west end of the glacier korridoren. the climb took them 17 days to complete, after which they were lifted by helicopter back to constable pynt. [climb magazine news letter, january 2009.] 2007 north liverpool land expedition: jimi gregson jim gregson led a party of six that visited north liverpool land from 7 to 21 april. from constable pynt they travelled by skidoo to the head of carlsberg fjord. a total of seven climbs were made, the highest 770 m. [climb magazine newsletter, january 2009.] 2007 n–s traverse of liverpool land: phil poole phil poole led a party on a north to south ski traverse of liverpool land. they travelled with jimi greg son’s skidoo party to carlsberg fjord, from where the traverse was to begin. the traverse was successful, with the last stage to ittoqqortoormiit being completed by dog sledge. both phil poole’s party and james gregson’s group flew back to europe from constable pynt on 21 april. [climb magazine newsletter, janu ary 2009.] 2007 south liverpool land: eduard birnbacher a german climber, eduard birnbacher, travelled about 15 km north of scoresbysund/illoqortoormiut 113 (ittoqqortoormiit) with a greenlandic assistant, and made two solo ascents between 13 and 23 april; the north-east pillar of kronen and an 800 m high summit south of kronen. [climb magazine news letter, january 2009.] 2007 grea sagax-revo and ecopolaris expeditions to east greenland these expeditions were organised by groupe de re cherches en écologie arctique (grea – arctic ecology research group). the sagax-revo party car ried out ivory gull censuses and botanical sampling on the ice cap south of scoresby sund (69°45´n 28°23´w) in june 2007, before flying to station nord to carry out further studies. the ecopolaris group carried out studies around holm bugt on traill ø in july and august, a continuation of a long-running grea research project (see grea 2003). visits were also made to areas on ymer ø and around forsblad fjord. [dpc report archive.] 2007 east greenland sortebræ expedition: david jakulis the first group of this 8-person expedition flew by twin otter via constable pynt arriving in the area west of sortebræ (69°01´n 27°51´w) on 9 june. the plane buried its nose in soft snow on landing, and took some time to dig out, helped by extra personnel landed by helicopter. the second party was flown out on 11 june, but due to snow conditions was landed some 40 km away from the first group. despite these problems, the two groups were reunited and attempted or climbed a number of peaks. four of these were north of 69°n latitude. the party flew back to iceland on 29 june. [bmc report archive.] 2007 ogwen valley dronning louise land expedition: russ hoar a three-person expedition comprising members of the ogwen valley mountain rescue organisation visited the nunatak region of south-west dronning louise land in may–june. travel to and from the region was by twin otter. the constant strong katabatic winds were a problem, as it was bitterly cold. numerous minor nunataks from 1900 to 2240 m high were climbed and claimed as first ascents. however, as the ice-cap surface is at c. 1800 m, none of the climbs involved ascents of more than a few hundred metres, and none were difficult. [bmc & rgs report archives.] 2008 arctic summits expedition in april and may 2008 georg czak and dominik rind made a long ski journey to the watkins bjerge, and via the gronau nunatakker to paul stern land. a total of six first ascents were made, as well as climbs of the four highest mountains in greenland (south of 69°n) around and including gunnbjørn fjeld. they had been set out by helicopter high on the ice cap and were picked up by twin otter from paul stern land. [dpc report archive.] 2008 paul stern land three british climbers (geoff bonney, jim and sandy gregson) were dropped off by twin otter in southwest paul stern land on 24 may. from their landing point they moved to northern paul stern land where they set up a base camp at 1800 m. the group suffered from strong katabatic winds but five first ascents were made, the highest being ararat. on 7 july the party was picked up by by a twin otter bringing in a nigel edwards climbing group (see below). [climb maga zine newsletter, january 2010.] 2008 nunataks north of paul stern land: nigel edwards nigel edwards led a six-person group of climbers that explored the nunatak region north of paul stern land. over the next 2½ weeks a total of 11 first and second ascents were made. the rock was reported as very poor, and none of the ascents involved more than about 500 m of vertical gain. [climb magazine news letter, january 2010.] 2008 greenland renland expedition: nat spring a three-member british expedition led by nat spring visited renland in june and july. the party flew into constable pynt airfield and on 27 june was lifted by helicopter to their base camp established on the lower part of edward bailey gletscher. three new peaks were climbed, and in the course of the expedition the party travelled the full length of edward bailey gletscher. a helicopter lifted them back to constable pynt on 21 july. [bmc & rgs report archives.] 114 2008 queens university belfast mountaineering club expedition: anthony garvey this six-person expedition led by anthony garvey visited renland in june. from constable pynt the party was lifted by two helicopter flights to a base camp set up on edward bailey gletscher. snow conditions were worse than in 2007 (west lancashire scouts expedition), and planned climbing routes had to be modified. the party climbed two summits by ski, and two impressive peaks on rock and ice. return to constable pynt was made by helicopter in a single flight. [bmc & rgs report archives.] 2008 the aktiv celebrates the centenary of the 1906–08 danmark-ekspeditionen the ice-strengthened wooden schooner aktiv visited northern east greenland as part of the centenary commemoration of the 1906–08 danmark-ekspedi tionen. the ship provided a working platform for geological investigations by geus (geological sur vey of denmark and greenland) in connection with the 2008 international polar year. leaving copen hagen on 3 july, the aktiv carried out a number of geological tasks and called at a several historical localities, including hekla havn, ella ø, mestersvig, daneborg and kap sussi, and arrived at danmarks havn on 23 august. the ship arrived back in copen hagen on 10 september (n. mikkelsen 2009). 2008 expedition blosseville kyst: pascal hémon a six-person expedition led by pascal hémon visited the blosseville kyst in july–august, using the 16 m aluminum yacht mio palmo for transport. a trip was also made along the liverpool land coast to the mouth of kong oscar fjord. an attempt was made to photograph rigny bjerg from the sea, first seen during the 1833 voyage of jules de blosseville, but poor weather hindered observations. [dpc report archive.] 2008 dresden university r/v polarstern voyage mirko scheinert led a four-man team aboard the r/v polarstern with the objective of setting up new gps stations at ice-free locations on the northern east greenland coast. in june and july a total of 16 new stations were established between 74° and 81°n, and 10 stations of the kms (kort& matrikelstyrelsen: national survey and cadastra) geodetic network were re-observed. [dpc report archive.] 2008 odder museum and danish arktisk institut eight curators from odder museum and danish arktisk institut joined 48 paying participants aboard the russian cruise ship aleksey marishev for a maritime archaeological cruise organised by the travel company oceanwide. this voyage in september 2008 was part of a series of initiatives to commemorate the 100th anniversary of the 1906–08 dan markekspeditionen. in danmark havn part of the telephone cable that had connected the ship danmark with the buildings on shore in 1906–08 was observed, as well as an abandoned iron pot. at snenæs a search was made for the motor vehicle that sank through the ice here in 1907, but no trace was found. off the north-east coast of shannon a search was made for the wreck of the alabama that had sunk off alabama havn in march 1910, but only the anchor was found. [dpc report archive.] 2008 ‘norfra’ winter expedition to north-east greenland: hans lapstun this three-person expedition led by the norwegian hans lapstun visited the region around nyhavn and mestersvig airfield in april and may. a month was spent at washburn’s hus, and for the last few weeks nyhavn was used as a base. [dpc report archive.] 2008 kayak expedition daneborg to ella ø two greenlanders from aasiaat in west greenland made a kayak tour through the fjords of northern east greenland between 28 july and 30 august. due to ice conditions their route was from daneborg, west of clavering ø, through loch fyne, a portage to the head of muskusoksefjord, and via ymer ø to ella ø. transport to daneborg and from ella ø was with twin otter. [dpc report archive.] 2008 the professor molchanov visit to the north-east greenland national park a group from ‘foreningen af danske biologer’ (so 115 ciety of danish biologists) took advantage of a cruise by the professor molchanov to visit sites in the north-east greenland national park that relate to danish-norwegian expeditions and whaling activities. between 8 and 11 september the group visited dan markshavn, germaniahavn, ella ø, alpefjord and scoresbysund. [dpc report archive.] 2008 eastern liverpool land simon burke and olly sanders undertook a kayak tour along the outer coast of southern liverpool land, and made stops at raffles ø, rathbone ø and kap høegh. some short climbs were made, and the party experienced problems with curious polar bears on several occasions. [climb magazine newsletter, january 2010.] 2008 kayak expedition mestersvig to ella ø: morten asklund a four-man expedition led by morten asklund made a problem-free kayak trip from mestersvig to ella ø, and return, between 28 july and 12 august. [dpc re port archive.] 116 117 a. schmidt gletscher 74ø-161 (74°01.8´n 22°26.1´w). minor glacier in the nørlund alper, ne hudson land, draining north to wordie bugt, so named by lauge koch’s 1929–30 expeditions. lacmann’s (1937) maps apply this name to nippoldt gletscher, immediately to the west. a. stelling sund 76ø-47 (76°22.0´n 20°28.0´w; map 4). sound between djævleøen and nanok ø. named by the 1906–08 danmark-ekspeditionen for anton stelling, who supplied paint to the expedition from his business in copenhagen without charge (j. løve, personal communication 2009). (a. stellings sund, stellings sund.) a.b. drachmann gletscher 76ø-131a (76°10.0´n 24°27.0´w; maps 2, 4; fig. 21). large e–w-trending glacier in carlsbergfondet land, southern dronning louise land. the area was mapped by lauge koch on reconnaissance flights in 1932 during the 1931–34 treårsekspeditionen, and the name was originally applied to the present budolfi isstrøm farther to the north. however, since budolfi isstrøm had been named by j.p. koch in 1917, the name a.b. drachmann gletscher was transferred to the present glacier. the name commemorates anders bjørn drach mann [1860–1935], professor at the university of copenhagen 1905–26, and chairman of the board of the carlsbergfondet 1926–33. (a.b. drachmann bræ.) a.p. olsen land 74ø-181 (74°38.0´n 21°40.0´w; maps 2, 4). land area between svejstrup dal and inner tyrolerfjord. mapped by lauge koch on reconnaissance flights in 1932 during the 1931–34 treårsekspeditionen (fig. 15), and named after anders peter olsen [1862–1932], colony manager in jakobs havn [illulissat] 1902–12 and later head of a department in grøn lands styrelse. (a.p. olsens land.) in this section all officially approved, and many unapproved, names are listed, together with explanations where known. approved names are listed in normal type or bold type, whereas unapproved names are always given in italics. names of ships are given in small capitals. individual name entries are listed in danish alphabetical order, such that names beginning with the danish letters æ, ø and å come after z. this means that danish names beginning with å or aa (e.g. aage bertelsen gletscher, aage de lemos dal, åkerblom ø, ålborg fjord etc) are found towards the end of this catalogue. å replaced aa in danish spelling for most purposes in 1948, but aa is commonly retained in personal names, and is optional in some danish town names (e.g. ålborg or aalborg are both correct). however, greenlandic names be ginning with aa following the spelling reform dating from 1973 (a long vowel sound rather than short) are treated as two consecutive ‘a’s. in the reference list of this volume the standard english alphabetical order is used. in each individual name entry the name (e.g. a. schmidt gletscher) is followed by the place name committee reference number (e.g. 74ø-161) and then the latitude and longitude in degrees, minutes and tenths of a minute (e.g. 74°01.8´n 22°26.1´w). alternative approved names are given in square brackets. description and explanation of the name then follows, and each entry closes with any recorded name variations in italics. greenlandic names are spelt according to the modern greenland orthography (spelling reform 1973), with cross-references from the old-style spelling still to be found on many published maps. prospectors place names used only in confidential company reports are not found in this volume. in general, only selected unapproved names introduced by scientific or climbing expeditions are included. incomplete documentation of climbing activities by expeditions claiming ‘first ascents’ on milne land and in nunatak regions such as dronning louise land, has led to a decision to exclude them. many recent expeditions to dronning louise land, and other nunatak areas, have gained access to their region of interest using twin otter aircraft, such that the remaining ‘climb’ to the summits of some peaks may be as little as a few hundred metres; this raises the question of what constitutes an ‘ascent’? an exception is made for climbs in the stauning alper (map 5), where there is almost full documentation by visiting expeditions with many climbing reports either published, or deposited in the report archives of the danish polar center (dpc), the royal geographical society of london (rgs) or british mountaineering council (bmc). in this section north-east, north-west, south-east and south-west are replaced by ne, nw, se and sw. catalogue of place names in northern east greenland a 118 aamarsuit [ikkaalissat] 70ø-294 (70°27.7´n 22°14.5´w). abandoned coal-mine, a small quarry on the coast of sw liverpool land east of aamarsuit nuaat. recorded by the 1955 geodætisk institut name registration, the name means ‘coal’. aumarssuit (= aamarssuit) was said in 1955 to be the name used by the younger generation. it has also been called dortes kulmine. see also ikkaalisaat. (aumarssuit). aamarsuit nuaat [basaltnæs] 70ø-293 (70°27.4´n 22°16.1´w). minor cape east of ittaajimmit [kap hope], sw liverpool land. recorded by the 1955 geodætisk institut name registration, the name translates as ‘coal cape’, a reference to nearby outcrops of coal. (aumarssuit nûat, íkauligssat nûat.) aantuuntap taartaa 70ø-297 (70°28´n 22°13´w). large stone on the west side of rosenvinge bugt, southernmost liverpool land. the name was recorded by the 1955 geodætisk institut name registration, and translates roughly as ‘antonies helping spirit’. it recalls an incident when henrik høegh’s kivfak (house-keeper) was said to have seen a spirit-being at this point. (ântûntap târtâ.) aappalaartukajik [røde hytte] 70ø (70°32´n 23°41´w). name used in grønlands landsmuseum reports for the inuit ruins around the present hunting hut known as røde hytte in southern jameson land. it translates as ‘the little reddish’. aappalaatsiaq 73ø-387 (73°36.7´n 25°31.4´w). this is probably the 1998 m high summit on the north side of grejsdalen in andrée land. it was named by erdhardt fränkl during lauge koch’s 1948–50 expeditions, although the position is not found on any of his maps (fränkl 1953). it is greenlandic for ‘the red’, and the name derives from the colour of the rocks. (augpalâtsiaq, apalatiak.) aappaleqisaap kuua [tværelv] 70ø-187 (70°31.5´n 22°10.2´w). river in southernmost liverpool land, draining into the west side of hvalrosbugt. recorded by the 1955 geodætisk institut name registration, the name translates as ‘reddish river’. (augpaleqisâp kûa.) aappaleqisaaq 70ø-298 (70°30.0´n 22°12.0´w). hillside on the west side of rosenvinge bugt, southern liverpool land, corresponding to the flanks of gulefjeld (= yellow hill). one of the names recorded by the 1955 geodætisk institut name registration, it is named for its colour, translating as ‘the reddish’. (aug pa le qisâq.) aappaleqisaaq kiatteq 70ø-189 (70°30.7´n 22°06.5´w). delta or slope on the west side of hvalrosbugt, southern liverpool land, across which aappaleqisaap kuaa [tværelv] drains. recorded by the 1955 geodætisk institut name registration, the name translates roughly as ‘east of the reddish’. (augpaleqisâq kiáteq.) aaronip sarpaa 72ø-284 (72°14.0´n 23°46.5´w; see also fig. 66). narrow entrance channel to noret, the enclosed bay near to mesters vig airfield. the name was recorded by the 1955 geodætisk institut name registration, and translates as ‘aron’s current’, a reference to strong tidal currents. (âronip sarpâ.) abraxas 72ø (72°05.6´n 25°12.3´w). peak 1900 m high on the south side of gully gletscher where it meets cavendish gletscher, stau ning alper. it was named by the 1984 paternò expedition, which made the first ascent on 1 august 1984. ‘abraxas’ is a mystic word found engraved and sometimes personified as a half-animal halfhuman deity on gemstones used as charms up to the 13th century. absalon havn 76ø-79 (76°39.8´n 18°50.0´w). small bay on the ne coast of store koldewey. named by the 1906–08 danmark-ekspe ditionen as absalons havn, possibly for the schooner absalon, a 341-ton trainingship of the danish navy, in its turn named after bishop absalon [1128–1201], founder of the city of copenhagen. achnacarry spids 72ø-363 (72°10.6´n 24°51.2´w; map 5). peak about 2130 m high in the north stauning alper at the head of dunottar gletscher. first climbed and so named by malcolm slesser’s 1958 expedition for achnacarry castle, inverness, home of clan cameron, which was burnt down in 1746. achnacarry house now stands on the site. (achnacarry.) achton friis ø 78ø-40 79ø-28 (78°57.6´n 19°13.6´w; map 4). island north of schnauder ø, jøkelbugten. named by eigil knuth’s 1938–39 mørkefjord expedition for johannes achton friis [1871–1939], a danish artist and writer. he was an artist on the 1906–08 danmark-ekspeditionen, when he made about 100 paintings and drawings. ad astra iskappe 77ø-137 (77°00.0´n 24°00.0´w; maps 2, 4; fig. 21). ice cap in northern dronning louise land, east of the lower part of admiralty gletscher. named by the 1952–54 british north greenland expedition as adastra iskappe in honour of the royal air force which supplied transport to the expedition, and whose motto is ‘per ardua ad astra’ (through difficulties to the stars). the current approved form was retained despite efforts by brian roberts on behalf of the expedition to change it to adastra iskappe. see also adastra lake. ad. s. jensen land 75ø-41 76ø-345 (76°06.0´n 21°08.0´w; maps 2, 4). land area north of bessel fjord. one of the names found on the 1932 edition of the geodætisk institut 1:1 million scale map, it derives from lauge koch’s aerial observations during the 1931–34 treårsekspeditionen. it was named after adolf severin jensen [1866–1953], a zoologist noted for his fishery investigations in west greenland, and professor at the university of copenhagen 1917–37; he was a member of the committee of the 1931–34 treårsekspeditionen. adam af bremen dal [william smith dal] 72ø-173 (72°48.8´n 22°31.2´w; map 4). e–w-trending major valley on se geogra phical society ø. the name was one of a group given by the place name committee in 1939 to replace proposals by hans stauber. adam af bremen [d. 1075], is noted for ‘de hamburgske ærke biskoppers historie’, a description of scandinavia based on written and spoken sources in which greenland is described. the valley is more usually known by its second authorised name, william smith dal. it was also called brandal by norwegian scientists. adastra lake 77ø (77°03.5´n 23°05.0´w). ice-dammed lake in eastern dronning louise land, which periodically develops on the site of southern strandelv. it was present in 1951 during the british north greenland reconnaissance expedition, when it was surveyed as a possible landing site for sunderland aircraft. see also ad astra iskappe. admiralty gletscher 76ø-309 77ø-131a (77°04.0´n 24°14.0´w; maps 2, 4; fig. 21). glacier in northern dronning louise land draining from the inland ice into støvdal. the name was given by the 1952–54 british north greenland expedition to commemorate the help given to the expedition by the royal navy. several of the expedition members were from the royal navy, and the admiralty also made available a secretary and the expedition headquarters in london. admiralty lake 77ø (77°08.6´n 23°24.6´w). name given to bri t annia sø in north dronning louise land by the 1951 british north greenland reconnaissance expedition, but changed to britannia sø when it became the site of the 1952–54 expedition base (banks 1957). slamsøen has also been used. adolf hoel gletscher 73ø-579 74ø-384a (74°00.0´n 27°30.0´w; maps 2, 4). name used for the e–w-trending glacier south of arnold escher land by the 1931 høygaard and mehren expedition, originally in the form adolf hoels bre. the name is now used in a more restricted sense than the original, and is confined to the ne–sw-trending part of the glacier. adolf hoel, a norwegian geologist and director of nsiu (see also hoelsbo), had provided transport for the expedition, and wrote the preface to the expedition narrative (høygaard & mehren 1931). adolf jensen bjerg 79ø-33 (79°46.4´n 19°44.4´w). mountain on sw hovgaard ø, nw of kap adolf jensen. named by john haller during lauge koch’s 1956–58 expeditions. see also ad. s. jensen land. afgrunden 73ø-63 (73°41.0´n 22°38.9´w). valley in hudson land 119 west of stordal. so named during lauge koch’s 1929–30 expeditions in the form afgrund valley, because it is a hanging valley with a cliff (= afgrund) at its mouth. agardh bjerg 73ø-519 (73°45.2´n 25°30.0´w). mountain 1820 m high in ne andreé land, on the west side of geologfjord. it was named by a.g. nathorst’s 1899 expedition as agardhs berg, probably for the swedish botanist jacob georg agardh [1813–1901], professor of botany at the university of lund from 1847. nathorst was at the university of lund from 1868–71, where he had originally intended to study botany, although his interests subsequently became palaeobotanical. (mount agardh, agardhs plateau). agardhskløft 73ø (73°46.6´n 25°32.8´w). name used during lauge koch’s 1948–50 expeditions for a ravine nw of agardh bjerg, andrée land (e.g. fränkl 1953). agassiz bjerg 73ø-332 (73°29.0´n 22°29.1´w). mountain on central gauss halvø. named during lauge koch’s 1936–38 expeditions by wolf maync and andreas vischer for louis agassiz [1807–73], a swiss palaeontologist. agassiz was noted especially for his studies of living and fossil fishes, and for his theories of widespread glaciation. agassiz dal 72ø-428 (72°55.5´n 27°42.8´w; map 4). valley in southern goodenough land, named during the 1931–34 treårs ekspeditionen by eugène wegmann in the form agassiz valley. see also agassiz bjerg. agatdalen 69ø (69°54.0´n 23°56.3.0´w). cleft leading up to the plateau on the sw side of steward ø where tuborg & sandell (1999) reported finding mounds of loose agate blocks, interpreted as raw material mined by the inuit for use as tools and weapons. agda dal 73ø-284 (73°22.1´n 23°04.0´w). valley on the sw coast of gauss halvø. named during the 1931–34 treårsekspeditionen by gunnar säve-söderbergh as agda valley, after agda brasch, a technical assistant at the riksmuseum, stockholm. aggersborg 72ø-225 (72°02.9´n 23°56.5´w; map 5). mountain south of mesters vig. named by prospecting teams associated with lauge koch’s 1948–49 expeditions after the viking fortress of the same name near aggersund, jylland, denmark. agnes-tufta 74ø (c. 74°40´n 20°13´w). norwegian hunting hut 3 km west of kap schumacher on the west side of albrecht bugt, northern wollaston forland. it was built in september 1928 by the hird expedition, and named after the youngest daughter of jørgen furnes who had helped build the hut; she was born after jørgen furnes left for greenland in the summer of 1927. the hut was moved in august 1930 to southern kuhn ø where it was known as furnes. agnesbjerg 74ø-407 (74°02.7´n 22°34.5´w). mountain in the nørlund alper, ne hudson land. this name appears to have been suggested by the place name committee in 1963 as a substitute for a proposal by paul stern. girl’s name. agnete sø 75ø-44 (75°38.4´n 20°16.8´w; map 4). lake in nw hochstetter forland. the lake was visited by norwegian hunters and danish scientists in l933, and the name subsequently appeared as a reference locality in botanical reports of the 1931–34 treårs ekspeditionen. girl’s name. a norwegian hut at the south point of the lake is known as agnetehytten. another hut, said by some sources to have been built in 1948 and to be known as agnete sø, was never erected according to p.s. mikkelsen (1994, 2008). (agnete lake, agnetes sø.) agnetehytten 75ø (75°35.6´n 20°00.1´w). norwegian hunting hut built by john giæver’s expedition in the spring of 1933 at the south point of agnete sø, nw hochstetter forland. (agnete sø hytten, agnetesøen.) agnetenæs 74ø (74°09.0´n 20°25.6´w). minor cape on the se coast of clavering ø, just east of basaltkap. the name appears on a sketch map in gustav thostrup’s 1921 logbook (in: møller 1939). girl’s name. agnetesøelven 75ø-91 (75°35.4´n 19°55.8´w). river draining ag nete sø, northern hochstetter forland. named by the danish hun ting company nanok in 1939. agpaliarsseqarteq – see appaliarseqarteq. agpaliarssoqarfik – see appaliarsoqarfik. agpalik – see appalik. agpalîp timâ, agpalîp tunua – see appaliip timaa, appaliip tunua. agsutsund – see assutsund. agurkesø 76ø-350 (76°47.0´n 18°39.6´w). lake close to dan markshavn weather station. the name was reported by hans meltofte as in general use by the staff at the weather station in 1969–71, and derives from the shape of the lake (agurk = cucumber); it was later approved with other meltoft suggestions. aiguille de jardin 71ø (71°55´n 25°52´w). mountain on the west side of prinsessegletscher, south of furesø. named and first climbed by the 1968 claude rey expedition. origin of name uncertain. ailsa 75ø-2 (75°18.0´n 19°37.3´w; map 4). hill 196 m high in hochstetter forland. named by douglas clavering in 1823 for its resemblance to the island ailsa craig in the firth of clyde. scottish maps of the early 19th century used the short form ‘ailsa’. (ailsa bjerg, mt. ailsa, ailsa hill.) ailsa sø 75ø-111 (75°17.8´n 19°44.5´w). small lake west of ailsa, hochstetter forland. the name was first used by the 1976 swedish-danish expedition which core-sampled the sediments in the lake (see also björck et al. 1994). ailsahytten 75ø-101 (75°17.0´n 19°22.5´w). danish hunting hut on the east coast of hochstetter forland, east of ailsa, built by nanok in august 1933. aina dal 73ø-115 (73°23.8´n 23°11.4´w). ravine on the sw coast of gauss halvø. so named by gunnar säve-söderbergh during the 1931–34 treårsekspeditionen, for aina stensiö, wife of erik a:son stensiö. see also stensiö bjerg. ainadalhytten 73ø (73°23.0´n 23°11.6´w). name often used for the norwegian hut built in october 1930 a few kilometres east of the mouth of aina dal, gauss halvø. it has also been known as von krogh or krogh-hytta. ajorpoq 73ø-395 (73°33.4´n 25°02.1´w). mountain in eastern andrée land, ne of the mouth of grejsdalen. so named by erdhart fränkl during lauge koch’s 1948–50 expeditions because his greenlandic assistant did not like it. the name translates as something ‘bad’. (agiorpoq.) ajúngilaq – see ajunngilaq. ajunngilaq 73ø-580 (74°00.9´n 28°57.4´w). nunatak 2284 m high in southern hobbs land. the name was used by arne høygaard and martin mehren in 1931 in the form ajungilakfjellet, and was employed in a broader sense than the present to include the whole of the present hobbs land. the nunatak seemed initially to threaten their progress, but was found to mark the western extension of the flat and easily negotiable adolf hoel gletscher. ‘ajungilak’, an inuit word for something ‘very good’, became their motto and is the title of the expedition narrative (høygaard & mehren 1931). the nunatak was climbed by hans katz on 8 august 1951. (ajúng ilaq, ajungilak.) akileqitâ – see akileqitaa. akileqitaa 69ø-63 (69°37.4´n 23°33.3´w). narrow cape or point in east henry land, on the northern blosseville kyst. the name was recorded by the 1955 geodætisk institut name registration, and translates as ‘that lying inbetween’. (akileqitâ, aqiliqitâ.) akínartequtâ – see akinnarteqitaa. akinnarteqitaa 69ø-56 (69°54.0´n 23°06.0´w; map 4). peninsula west of steward ø, northern blosseville kyst. one of the names recorded by the 1955 geodætisk institut name registration, it derives from its position between two fjords, translating as ‘that lying inbetween’. (akínartequtâ.) akselborg 72ø-249 (72°17.1´n 24°27.9´w; map 5). mountain in the northern stauning alper, sw of syltoppene. named by erdhart 120 fränkl during lauge koch’s 1950–51 expeditions, originally in the form gammel axels tinde (fränkl 1953), after axel jensen, skipper of the polypen in 1950–51. the name was altered to the present form by the place name committee, apparently to disguise the fact that it was named after a living person. akuliaruseq janet watson 76ø (76°28.3´n 22°26.1´w). peninsula at the head of bræfjorden, west of dove bugt. the name was proposed by brian chadwick, following his geological mapping in the region with the 1988–90 ggu north–east greenland project. it commemorates janet v. watson [1924–85], an eminent british geologist noted for her contributions to the understanding of the evolution of complex remobilised gneiss terrains. (janet watson halvø.) alabama 75ø-70 (75°17.2´n 17°50.5´w; map 4; fig. 26). hut in ne shannon built from the timbers of the alabama, a 50-ton sloop purchased and strengthened for the 1909–12 alabama expedition. the hut was used by østgrønlandske fangstkompagni from 1920 to 24, and from 1929 was taken over by nanok. the hut is still standing, surrounded by a variety of debris salvaged from the wreck of the alabama, but even in mid-summer contains icy snowdrifts inside and is uninhabitable (1988). (alabamahytten, ala ba ma hus, alabama hus.) alabama havn 75ø (75°17.2´n 17°49.8´w). small bay on the east coast of shannon adjacent to the hut alabama. the alabama wintered here in 1909–10, and sank in the bay in march 1910. see also alabama. (alabamahavn.) alabama nunatak 77ø-52 (77°44.6´n 23°53.2´w; maps 1, 2, 4). nunatak west of hertugen af orléans land, so named by the 1909–12 alabama expedition for the expedition ship. see also alabama. alabamablick 75ø (c. 75°19´n 17°48´w). feature in the vicinity of the base camp of the 1943–44 german meteorological station at kap sussi, shannon. the name is recorded by olsen (1965). it apparently had a view to the south of the hut alabama. albert 71ø (71°47.1´n 25°30.7´w; map 5). peak about 2300 m high in the southern stauning alper between borgbjerg gletscher and orion gletscher. named and climbed by the 1971 university of lancaster expedition. albert heim bjerge 74ø-326 (74°04.9´n 23°12.6´w; map 4). mountain range on the north side of promenadedal, south of wordie gletscher. so named during lauge koch’s 1936–38 expeditions by heinrich bütler for one of the most noted of swiss geologists, albert heim [1849–1937]. he was a structural geologist and professor at the university of zurich 1875–1911, and was celebrated for his studies of alpine geology. (albert heimberge, albert heims bjærge.) albrecht bugt 74ø-36 (74°36.0´n 19°47.0´w; fig. 15). large bay on the north coast of wollaston forland. named by karl koldewey’s 1869–70 expedition as albrecht bai (fig. 6) for george alexander albrecht [1834–98], treasurer of the ‘bremisches comité für die zweite deutsche nordpolarfahrt’, one of the principal expedition support organisations. (albrechts bugt, albrecht bay, albrecht bugten.) albrechts-slette 74ø (74°34.0´n 19°58.0´w). name used by danish hunters and others (e.g. christensen 1965) for the plain in northern wollaston forland, sw of albrecht bugt, which is officially known as storsletten. (albrechtsletten.) albrechtbugthytten 74ø (74°35.7´n 19°51.4´w). sirius hut built in august 1960 about 2 km nw of the head of albrecht bugt, northern wollaston forland, adjacent to the old norwegian hut (sletta) built in august 1928 by the hird expedition and known as albrechtsbugthytten. albuen [nuugaatsaa] 70ø-144 (70°34.4´n 22°34.7´w). cape on the west side of hurry inlet, so named by alfred rosenkrantz during lauge koch’s 1926–27 expeditions for its shape (albuen = the elbow). aldebaran gletscher 71ø-285 (71°53.8´n 24°08.4´w; map 5). glacier in the sw werner bjerge flowing west to join schuchert gletscher. the name first appeared on the maps of styger (1951), in his report on a climbing excursion during lauge koch’s 1950 expedition. it was named after the red giant star aldebaran (= the follower) in the constellation taurus. several other features in the region were named after constellations or planets. aldersro – see kap helgoland hytten. aldinger elv 70ø (70°40.0´n 25°35.7´w). major south-flowing river on se milne land. the name appears on the maps of callomon & birkelund (1980) and larsen et al. (2003), and commemorates hermann aldinger, a geologist who made pioneer studies in the region in 1933. attempts to obtain approval of the name in 1977 failed on the grounds that he was then still alive. aletschhorn 73ø-682 (73°36.3´n 27°24.9´w; map 4). mountain in eastern louise boyd land, west of gerard de geer gletscher. it was named by john haller during lauge koch’s 1949–51 expeditions, fig. 26. alabama, the hut on north-east shannon built from timbers rescued from the ship alabama that sank in its winter harbour nearby in march 1910. the hut is surrounded by a variety of debris from the ship, including a large rusty ice-saw in the left foreground. 121 after the mountain of the same name in central switzerland. alf bruns red 76ø (76°03.2´n 20°04.5´w). anchorage off bessel fjord hunting station, north of the mouth of bessel fjord near vesternæsset. so named by the 1932 gefion expedition, which anchored here, after captain alf brun [1866–1932], one of the committee of østgrønlandske fangstkompagni. (alf bruuns red.) alfabet nunatakker 71ø-380 (71°54.0´n 30°05.5´w; maps 3, 4). group of nunataks in western charcot land, extending from beta nunatak in the south to latitude 72°n. during geological mapping on the 1968 ggu expedition, the different nunataks were for convenience labelled alphabetically. beta nunatak is the largest. alfred escher land – see arnold escher land. alfred wegener bjerg 71ø (71°50.0´n 25°36.0´w; map 5). peak in the southern stauning alper, in the inner ne part of borgbjerg glet scher. probably first climbed and named by the 1977 schwäbische stauning alper expedition. aliertinde 72ø (72°07.3´n 24°58.5´w; map 5). rock peak on the sw ridge of dansketinden about 2580 m high. it was climbed and so named by the 1996 scottish mountaineering club expedition. allday dal 71ø-171 (71°43.9´n 23°22.7´w). valley draining north into ørsted dal, scoresby land. the name was one of a group given by the place name committee in 1939 to replace proposals by hans stauber. allday dal commemorates jacob allday, sent out by frederik ii of denmark in 1759 to rediscover greenland. allday hytte 71ø (71°45.6´n 23°23.8´w). norwegian hunting hut built by helge ingstad and normann andersen in 1932–33 in ørsted dal, at the mouth of allday dal, scoresby land. it was repaired by otto lapstun in 1982 as a memorial to norwegian hunting activities. the hut is also known as ørsted dal hytte. alliance col 71ø (71°50.4´n 25°20.0´w; map 5). high pass (2250 m) in the southern stauning alper between the upper bjørnbo gletscher (main glacier) and a branch of roslin gletscher. so named by the 1992 scottish stauning alper expedition for the ‘auld alliance’ between scotland and france, since the expedition included scottish and french members. alte hütte – see hansa bugt. alpebjerg 73ø-398 (73°28.0´n 25°32.0´w; map 4). mountain 2052 m high in se andrée land overlooking eleonore bugt. named during lauge koch’s 1948–50 expeditions by erdhart fränkl for its alpine character. alpedal 73ø-397 (73°28.0´n 25°27.5´w; map 4). valley in se andrée land draining into eleonore bugt, named by erdhart fränkl during lauge koch’s 1948–50 expeditions. alpedal 74ø-404 73ø-352 (74°00.7´n 25°23.0´w). valley in the central part of northern strindberg land. named during lauge koch’s 1948–49 expeditions by hans r. katz for its alpine character. alpefjord 72ø-27 (72°15.0´n 25°25.5´w; maps 3–5; fig. 27). n–strending fjord between nathorst land and the northern stauning alper. named alpfjorden by a.g. nathorst’s 1899 expedition for the spectacular high mountains of the stauning alper on the east side of the fjord. (alp fiord.) alpefjordhytten 72ø (72°17.4´n 25°20.5´w). norwegian hunting hut on the east side of alpefjord. it was built by helge ingstad’s expedition about 1932–33. (alpehuset.) alvinhögda 73ø (73°30.1´n 21°18.5´w). hill 365 m high in southern hold with hope, east of myggbukta. the name is found on an nsiu map (1932a), and is apparently derived from a personal name. alwin pedersens hus 76ø (76°55.1´n 20°06.5´w). hut built in august 1938 at hvalrosodden, adjacent to hvalrosodden station, and used by the zoologist alwin pedersen during the 1938–39 mørkefjord expedition. it was in good condition in 1990. amaroqarteq 71ø-201 (71°36.6´n 27°06.5´w). inuit ruin on the north coast of nordvestfjord, opposite the mouth of flyverfjord. recorded by the 1955 geodætisk institut name registration, the name translates as ‘where there are wolves’. ambolten 78ø-24 (78°18.2´n 19°13.6´w; maps 1, 4). island in jøkelbugten, named by eigil knuth’s 1938–39 mørkefjord expedition together with stigbøjlen and hammeren, for an apparent resemblance in shape to bones in the ear (ambolt = anvil). amdrup havn [ittoqqortoormiit kimmut kangertivat] 70ø-312 (70°28.4´n 21°54.5´w). small sheltered bay east of the settlement of scoresbysund [illoqqortoormiut], southern liverpool land. fig. 27. view south-east across alpefjord to the high summits of the stauning alper. from left: frihedstinde 2610 m, dansketinden 2842 m, norsketinden 2797 m, korsspids 2780 m and sefström tinde 2714 m. the john haller photograph collection, geus archive. dansketinden korsspidsfrihedstinde norsketinden alpefjord sefström tinde 122 first visited by otto nordenskjöld in 1900, the bay was named subsequently by the 1924–25 colonisation expedition after georg carl amdrup [1866–1947], a danish naval officer and greenland explorer. amdrup led the 1898–1900 carlsbergfondets expedition that in 1900 explored and mapped the east greenland coast from kap dalton (69°25´n) to agga ø (67°24´n). (amdrup harbour, amdrups-hafen, port amdrup.) amdrup hytte 69ø-18 (69°26.0´n 24°08.0´w). hut built by g.c. amdrup’s 1898–1900 expedition in a small bay on the north side of kap dalton, northern blosseville kyst. it was intended as an emergency wintering hut for the planned 1900 coast exploration, and features on expedition maps as amdrups depot. it was still standing in 1980, but reported to be in poor condition. (amdrups hytte.) amdrup land 80ø-10 81ø-128 (80°47.0´n 15°22.0´w; maps 1, 4). land area between antarctic bugt and ingolf fjord. it was named by the 1906–08 danmark-ekspeditionen after georg carl amdrup [1866–1947], a member of the expedition committee who had also presented the expedition with the boat used during his 1900 expedition. (amdrups land.) ameliebugt 77ø (77°25.0´n 19°15.0´w). name used by c.s. poulsen during the 1906–08 danmark-ekspeditionen for skærfjorden (lundbye 1984), the large fjord north of germania land which is bounded to the north by kap amélie. ammonitbjerg 70ø-117 (70°56.3´n 22°48.2´w). part of the sw slope of dusén bjerg, nw of the head of hurry inlet. named during g.c. amdrup’s 1898–1900 expedition, the name was first found in reports and on maps in the form ammonite mountain. madsen (1909) notes it as the locality where otto nordenskjöld had collected well-preserved ammonites in 1900. the name appeared on maps at the approximate position of the present eli bjerg for many years, until observations by tom harris (in: rosenkrantz 1934) showed that nordenskjöld’s mountain must lie farther inland. amoebites-elv 74ø (74°53.5´n 20°33.0´w). name used by wolf maync for a river on western kuhn ø. it derives from his work during lauge koch’s 1936–38 expeditions, and was given for finds of the fossil amoeboceras (maync 1947). (amöbites elv.) amphitheatre cliffs 77ø (77°34.8´n 21°03.3´w). cliffs west of dead lake in nordmarken. named by the 1987 irish expedition to north–east greenland. amphitheaterpingo 71ø (71°47.4´n 23°41.9´w). name used by fritz müller for the amphitheatre-shaped remains of a pingo in pingo dal, northern jameson land (müller 1959). amsjøhytten 75ø (75°16.1´n 21°25.2´w). norwegian hunting hut built in the spring of 1949 by arktisk næringsdrift about 14 km up in kildedal, c.h. ostenfeld land. the hut was built by eigil amsjø. an caisteal 72ø (72°03.5´n 24°59.9´w; map 5). mountain 2614 m high between the heads of gullygletscher and storgletscher, northern stauning alper. climbed and named by the 2007 smc east greenland expedition; the name means ‘the castle’. an dorus mor 71ø (71°47.3´n 25°30.3´w; map 5). name used by the 1996 norwegian stauning alper expedition for the pass between orion gletscher and borgbjerg gletscher, also known as the orionborgbjerg col. the name is derived from the gaelic and means the ‘great gate’. anden hvide 74ø-170 (74°21.5´n 20°37.4´w). part of a mountain range on ne clavering ø, named by arne noe-nygaard and gunnar säve-söderbergh during the 1931–34 treårs ekspe di tionen. originally three peaks were given the names erste weisse, zweite weisse and dritte weisse, for the colour of the rocks, names usually used in a shortened danicised form as 1. hvide, 2. hvide and 3. hvide (anden hvide = second white). see also første hvide and tredie hvide. anders jahre nunatak 73ø-573 (73°38.2´n 29°58.1´w; map 4). nunatak north of hamberg gletscher. so named by arne høy gaard and martin mehren in 1931 for anders jahre, a norwegian whaling-ship owner and lawyer. (anders jahres nunatak.) andreas lundager ø 76ø-212 (76°33.5´n 20°49.9´w). island in dove bugt north of godfred hansen ø. named by paul gelting during eigil knuth’s 1938–39 mørkefjord expedition for andreas lundager [1869–1940], the botanist of the 1906–08 danmarkekspeditionen. (lundagers ø, andreas lundagers ø.) andrée land 73ø-512 (73°40.0´n 26°17.0´w; maps 2, 3, 4). land area bounded by geologfjord, kejser franz joseph fjord, gerard de geer gletscher and adolf hoel gletscher. named by a.g. nat horst’s 1899 expedition for salomon august andrée [1854–1897], a swedish engineer who attempted to reach the north pole from spitsbergen by balloon in 1897 with two companions, but crashlanded on the ice and died on kvitøya (white island). one of the principal aims of nathorst’s 1899 expedition was to search for traces of andrée’s expedition. (andrées land.) andresensfjell 74ø (74°26.3´n 21°12.5´w). name used by norwegian hunters for a mountain on northern clavering ø, probably that which appeared on 1932 nsiu maps as tiedemannfjellet. it may have been named after herman andresen, who organised numerous hunting expeditions to the region. see also herman andresen fjellet. anduin 81ø (81°10.4´n 13°00.0´w). river draining se in east kilen, kronprins christian land. the name is found on a coloured geological map of kilen printed in 1991 (pedersen 1991), and was named after a locality in tolkien’s ‘lord of the rings’. angalassut nûat – see angalasut nuuat. angalasut nuuat 70ø-366 (70°29.2´n 21°58.7´w). cape to the west of scoresbysund [illoqqortormiut], southern liverpool land, probably identical with the original ferslew pynt. recorded by the 1955 geodætisk institut name registration, the name translates as ‘travellers cape’. the colonisation ship unloaded its cargo directly ashore at this point in 1924, and it was here that visitors to the settlement came ashore. (angalassut nûat.) angelin bjerg 73ø-528 (73°09.8´n 24°19.4´w). mountain 1900 m high on central ymer ø. a.g. nathorst’s 1899 expedition named it after nils peter angelin [1805–1876], a swedish palaeontologist and stratigrapher noted especially for his work in skåne, sweden. angelin had introduced nathorst to geology when he was a student at the university of lund. (angelin mountain, angelinfjellet.) anita ø 72ø-334 (72°40.8´n 22°42.2´w). small island in vega sund. the danish søkortarkivet proposed the name in 1956–57 when surveying the channel through vega sund as an alternative ap proach for ships on their way to nyhavn. it was named after the anita dan, a 3225-ton ice-strengthened polar ship built for the j. lauritzen shipping company for the greenland and finnish trade. in 1967 the ship was sold and rebuilt as the hms endurance, a british supply and ice-patrol vessel used in the antarctic. ankerbjerg 73ø-66 (73°36.3´n 22°33.7´w; map 4). mountain on the north side of moskusoksefjord. it was named by helge backlund during lauge koch’s 1929 expedition in the form mt. ankar for the anchorage on its south side. farther east moskus oksefjord becomes very shallow and unnavigable. (kap anker, ankerberg, ankar bg.) ankerbjergsdalen 73ø-723 (73°40.2´n 22°48.7´w). valley in south ern hudson land, in which ankerbjergselv flows, and which reaches the coast east of ankerbjerg. the name was approved at the suggestion of peter friend following his 1968–70 expeditions, although it had also been used occasionally earlier (e.g. backlund 1930). (ankar valley, ankerbergtal). ankerbjergselv 73ø-67 (73°40.2´n 22°48.7´w). river draining into moskusoksefjord east of ankerbjerg, named by lauge koch’s 1929–30 expeditions in the form anker river. ankerbukta 73ø (73°36.1´n 22°22.5´w). bay se of ankerbjerg in moskusoksefjord, an anchorage used by nsiu in 1929, and probably identical with ankerplads. (ankerhamna, ankerplassen). 123 ankerlien 73ø (73°36.5´n 22°28.5´w). norwegian hunting hut built in september 1929 by arktisk næringsdrift near ankerbjerg, about 6 km from the head of moskusoksefjord. it is also known as braasted. ankerplads 73ø (73°36.1´n 22°28.5´w). name used by gelting (1934) during the 1931–34 treårsekspeditionen for the anchorage se of ankerbjerg in moskusoksefjord. probably identical with ankerbukta. ankerpladsen 76ø-268 (c. 76°57´n 21°28´w). anchorage on the north side of inner mørkefjord, daniel bruun land, named by eigil knuth’s 1938–39 mørkefjord expedition which anchored the expedition motorboat here. ankervig 70ø-394 (70°21.3´n 28°09.5´w). small bay at the mouth of hjørnedal, on the north coast of gåseland. so named by the 1963 geodætisk institut expedition because it is possible to anchor small boats here. anna bistrup fjelde 79ø-34 (79°45.0´n 18°29.9´w; maps 1, 4). mountain on southern hovgaard ø, west of kap anna bistrup. named by john haller during lauge koch’s 1956–58 expeditions. see also kap anna bistrup. anna sten gletscher 74ø (74°50.0´n 22°22.4´w; fig. 15). name used for the present tvegegletscher, west of th. thomsen land. the name only appears on the 1932 1:1 million scale geodætisk institut map prepared by lauge koch during the 1931–34 treårs eks peditionen. it commemorates a russian-american film star whose first major success was in zola’s ‘nana’ in the usa. this was one of the few lauge koch name suggestions which the place name committee would not accept. annekssøen 77ø-24 (77°18.5´n 21°07.0´w; map 4). elongate nw– se-trending lake north of sælsøen. it was discovered by the 1906–08 danmark-ekspeditionen and named as annekssöen or annekssø. like nearby sælsøen it appears at one time to have been a fjord. (anneks sø, annexsø, anneks-søen.) annielva 73ø (73°28.0´n 21°17.8´w). stream in southern hold with hope, east of myggbukta. so named on the nsiu 1932 map (nsiu 1932a; fig. 13), probably for the anni 1, a norwegian sealer which brought the johan a. olsen expedition to greenland in 1922. the expedition founded the weather station at myggbukta, but was lost without trace when the anni 1 disappeared in 1923 on its way home through the coastal ice belt. ansgar 72ø-207 (72°10.3´n 23°59.2´w; map 4). one of the summits of korsbjerg, 1011 m high, south of mestersvig airfield. it was named by prospecting teams associated with lauge koch’s 1948–49 expeditions, probably after ansgar [801–865], arch bishop of hamburg. a missionary to denmark, he built the first danish church in slesvig in the year 850. annsketinde 72ø (72°06.6´n 24°58.5´w; map 5). peak 2460 m high on the southern spiky ridge of dansketinden. it was climbed and so named by the 1996 scottish mountaineering club expedition. antarctic bugt 80ø-14 81ø-129 (80°55.0´n 14°10.0´w; maps 1, 5). fjord or large bay on the north side of amdrup land. so named by the 1906–08 danmark-ekspeditionen, probably for the ship antarctic used on a.g. nathorst’s 1899 expedition to east greenland. the antarctic was a 353-ton, steam-driven bargue, built in drammen in 1871, which had sailed as a sealer and whaler under the name cap nor. renamed the antarctic in 1893 for a hunting trip to the antarctic, it was also used by nathorst for his expedition to spitsbergen in 1898 and to east greenland in 1899. in 1903 it was crushed by pack ice in the weddel sea during otto nordenskjöld’s antarctic expedition. antarctic dal 72ø (72°00.0´n 23°21.0´w). name used by bearth (1959) for the present kolledalen, which drains eastwards to antarctic havn, northern scoresby land. antarctic gletscher 71ø (71°57.9´n 23°49.4´w). name occasionally used by bearth (1959) for nedre antarctic gletscher in the werner bjerge, scoresby land. antarctic havn 72ø-19 (72°01.0´n 23°08.0´w; map 4). large bay on the south side of davy sund, northern scoresby land, so named antarctics hamn by a.g. nathorst’s 1899 expedition because the expedition ship antarctic anchored here on 20 august. see also antarctic bugt. tornøe (1944) suggested the harbour might correspond to the ‘finnsbúdir’ of the icelandic sagas. the hunting station at the head of the bay, originally known as karlsbak, has sometimes been referred to as antarctic havn station. (antarctic hamna, antarctic harbour.) antarctic havn station – see karlsbak. antarctic pas 71ø-248 (71°58.5´n 23°51.8´w; map 5). col on the east side of østre gletscher in the werner bjerge, scoresby land, leading east to kolledalen (sometimes called antarctic dal) and antarctic havn. named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk. antarctic spids 71ø-249 (71°58.8´n 23°53.0´w; map 5). mountain 1483 m high in the werner bjerge, scoresby land, north of antarctic pas. it was named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk, and climbed by bearth in 1953. antarctic sund 73ø-526 (73°07´n 25°30´w; map 4). sound con necting central kejser franz joseph fjord with kong oscar fjord. it was named by a.g. nathorst during his 1899 expedition for the ship antarctic. see also antarctic bugt. (antarctics sund, antarc tic strait, antarcticsundet, antarktiksundet.) antezedenzpingo 72ø (72°32.6´n 23°42.1´w). name given to a pingo in karupelv valley, traill ø, during lauge koch’s 1954–55 expeditions. so named because it developed across the former course of the river (müller 1959). anthons ø 76ø (76°43.8´n 20°32.0´w). island east of daniel bruun land, the present midterholmen. so named on maps of the l932 gefion expedition (jennov 1935), after anthon jensen, ship’s boy on the gefion. antiklinalbugt 72ø-278 (72°48.4´n 25°08.6´w; fig. 28). bay on sw ella ø, dominated by an imposing anticlinal structure in the rocks of the cliff behind the bay. named by john cowie during lauge koch’s 1949–54 expeditions. antoinette’s bjærg 74ø (74°25.3´n 19°51.1´w). name used by wolf maync during lauge koch’s 1936–38 expeditions for a mountain 992 m high in wollaston forland, north of hammeren (maync 1949). girl’s name. anton jensensundet 72ø (72°37.9´n 22°29.8´w). sound between nordenskiöld ø and kap palander in vega sund. used only on nsiu maps (lacmann 1937), and named after anton jensen [b. 1890]. as captain of the terningen, jensen sailed a norwegian hunting expedition to east greenland in 1928. antonsens hytte 74ø (74°30.9´n 21°10.7´w). norwegian hunting hut built in the summer of 1937 by gerhard antonsen for arktisk næringsdrift in store sødal, a.p. olsen land. ântûntap târtâ – see aantuuntap taartaa. apostlene 74ø-241 (74°29.0´n 18°59.0´w). name used by the 1908–09 floren expedition (brandal 1930) for two rock pinnacles on the mountain forming kap wynn, eastern wollaston forland (apostlene = the apostles). appaliarseqarteq 70ø-204 (70°32.4´n 21°29.2´w). cliffs between kap lister and kap hodgson, liverpool land. recorded by the 1955 geodætisk institut name registration, the name translates as ‘the place where there are little auks’. little auk colonies are found on many of the cliffs of eastern liverpool land. (agpaliarsse qarteq.) appaliarsoqarfik 70ø-351 (70°06.9´n 22°18.6´w). cliff west of kap brewster on volquaart boon kyst, forming the northern steep face of the mountain ridge known as søstrene. one of the names recorded by the 1955 geodætisk institut name registration, the name means ‘where there are little auks’. (agpaliarssoqarfik.) appaliip timaa 70ø-208 (70°32.9´n 21°33.7´w). coastal stretch of 124 the bay appaliip tunua, south liverpool land. recorded by the 1955 geodætisk institut name registration, the name means ‘the inner side of the place with little auks’. (agpalîp timâ.) appaliip tunua 70ø-207 (70°33.2´n 21°33.6´w). bay due south of appalik [raffles ø], se liverpool land. the name was recorded by the 1955 geodætisk institut name registration, and translates as ‘appalik’s back-side’. the local spelling has been recorded as appa lip dunua. (agpalîp tunua.) appalik [raffles ø] 70ø-209 (70°36.1´n 21°31.2´w). island in lillefjord, se liverpool land. one of the names recorded by the 1955 geodætisk institut name registration, the name means ‘here are little auks’. (agpalik.) appenzeller nunatak 72ø-464 (72°39.3´n 28°08.3´w). nunatak west of gletscherland in the upper reaches of hisinger gletscher. so named by eugène wegmann during the 1931–34 treårsekspe di tionen, for the inhabitants of the swiss canton of appenzeller, noted for their conservatism (behind the mountains and behind the times). the nunatak was reached by a geological party led by wegmann in august 1934. apuseeq [hvidefjeld] 70ø-202 (70°32.0´n 21°44.2´w). ice cap ne of scoresbysund [illoqqortormiut], south liverpool land. re corded by the 1955 geodætisk institut name registration, the name translates as ‘the big snow-patch’. (apusêq.) apusêq – see apuseeq. apusiaajiip nuaa 70ø-338 (70°26.5´n 21°40.0´w). cape ne of kap swainson, south liverpool land. recorded by the 1955 geodætisk institut name registration, the name means ‘apusiaajik’s cape’. (apusiâjîp nûa, apusiaajiip nuua.) apusiaajik 70ø-337 (70°26.7´n 21°41.2´w). area of perennial snow ne of kap swainson, south liverpool land. recorded by the 1955 geodætisk institut name registration, the name means ‘it has a lot of snow’. (apusiâjik.) apusiâjik – see apusiaajik. apusiâjîp nûa – see apusiaajiip nuaa. apusiikajik [aage nielsen gletscher] 70ø-215 (70°40.2´n 21°48.9´w). glacier in south liverpool land draining se into lillefjord. recorded by the 1955 geodætisk institut name registration, the name translates as ‘the rather large snow patch’. (apusî kajik.) apusîkajik – see apusiikajik. apusinikajik 71ø-215 (71°17.0´n 25°53.2´w). glacier at the head of skillebugt, east renland. recorded by the 1955 geodætisk institut name registration, the name means ‘the little snow’. aqartersiorpik 72ø-283 (72°11.2´n 24°07.2´w). lead mine, now abandoned, near mestersvig airfield. recorded by the 1955 geo dætisk institut name registration, the name translates as ‘where one looks for lead’. see also minebyen and store blydal. aqigssip kangertiva – see aqissip kangertiva. aqissip kangertiva [rypefjord] 70ø-3 71ø-38 (71°00.0´n 27°40.0´w). fjord between se renland and c. hofmann halvø. the name was recorded by the 1955 geodætisk institut name registration, and is a translation of the danish name (= ptarmigan fjord). (aqigssip kangertiva, aqissit kangersuat). arabertoppe 70ø-397 (70°36.7´n 26°18.0´w). mountain range in south milne land, north of rensund, named by the 1963 geo dætisk institut expedition for a supposed resemblance to a row of arab tents. ararat 70ø-435 (70°30.8´n 29°53.0´w). nunatak 2480 m high in nw paul stern land. named by laurent jemelin during the 1967–72 ggu scoresby sund expeditions after mount arrarat of the old testament, in association with the nearby nunatak known as arken. aravis 73ø-321 (73°51.9´n 22°47.0´w). mountain in hudson land north of dybendal. named during lauge koch’s 1936–38 expeditions by heinrich bütler for the mountain chain of the same name in the savoy alps. arbenz kolle 72ø-110 (72°43.1´n 25°20.3´w; map 4). domeshaped mountain in ne lyell land, with a summit ice cap, named during the 1931–34 treårsekspeditionen by eugène wegmann as arbenz-dome. paul arbenz [1880–1943] was a swiss structural geologist mainly known for his work in the alps. arbino bjerg 72ø-446 (72°46.9´n 27°13.3´w; map 4). mountain on the south side of dickson fjord, gletscherland. the name was used by eugène wegmann during the 1931–34 treårsekspedi tionen. archangel peak 72ø (72°04.5´n 25°05.4´w; map 5). summit 2558 m high on the west side of upper gullygletscher, northern stauning alper. climbed and named by the 2007 smc east greenland expedition. archer øer 72ø-21 (72°12.7´n 23°40.4´w; map 4). two islands on the south side of kong oscar fjord, named by a.g. nathorst in 1899 as archers öar for colin archer [1832–1930]. the son of a scottish ship builder who settled in norway, colin archer built the fram used by fridtjof nansen in his crossing of the arctic ocean, and also carried out improvements on the antarctic for nathorst. (archer island, archeröyane). arctic riviera 72ø, 73ø, 74ø (72°–74°n). popular name used for parts of northern east greenland with a generally favourable summer climate, but in particular that part of east greenland around ella ø where lauge koch’s geological expeditions were based. erik hofer’s ‘arctic riviera’ (1957) was illustrated largely by photographs taken during lauge koch’s summer expeditions. arcturus gletscher 71ø-261 (71°58.0´n 24°13.3´w; map 5). glacier in the werner bjerge flowing sw to join schuchert fig. 28. the bay on the west side of ella ø known as antiklinalbugt, named for the spectacular anticline developed in lower palaeozoic rocks. 125 gletscher. the name appeared first on the maps of styger (1951), a record of a climbing excursion during lauge koch’s 1950 expedition. it was named after the star arcturus. ardencaple fjord 75ø-4 (75°20.2´n 21°00.0´w; maps 2, 4; see also fig. 59). fjord between dronning margrethe ii land and c.h. ostenfeld land, which divides westwards into two branches known as bredefjord and smallefjord. it was named by douglas clavering in 1823 as ardencaple inlet for ardencaple castle, dum barton, the residence of his friend and relative lord john camp bell. (ardencaple-bai, ardencaple fiord.) ardvreck bjerg 72ø-353 (72°10.9´n 25°47.8´w). mountain 1866 m high in eastern nathorst land between sandgletscher and syd vestgletscher. it was first climbed by a malcolm slesser party in 1958, and named ardvreck for ardvreck castle, a macleod stronghold on the shore of loch assynt, sutherland, built in 1591 and now a ruin. arenaen 73ø-425 (73°19.5´n 24°46.9´w). small plateau on northern ymer ø, named by silvio eha during lauge koch’s 1947–49 expeditions (arenaen = the arena). arentzhytta 73ø (73°02.8´n 24°04.7´w). norwegian hunting hut on the north side of sofia sund, west of the mouth of barnabas dal and about 7 km east of rødebjerg. it was built in october 1929, and named after gustav a. arentz, a director of arktisk næringsdrift. it has also been known as snehytten and rødebjerghytten. (arentzhytten.) argand gletscher 72ø-105 (72°41.2´n 25°56.4´w). glacier in northern lyell land, draining north to kempe fjord. the name was used by eugène wegmann during the 1931–34 treårsekspedi tionen in the form argand glacier, and commemorates émile argand [1879–1940], a structural geologist especially known for his studies in the swiss alps. argandhornene 72ø-106 (72°42.4´n 25°49.3´w). mountain summits in northern lyell land, east of argand gletscher. named by eugène wegmann during the 1931–34 treårsekspeditionen. see also argand gletscher. (argandhorns.) aries glacier 71ø (71°35.0´n 25°00.0´w; map5). glacier in the southern stauning alper draining from first point of aries via gurreholm dal to schuchert dal. the name was first used by james clarkson’s 1961 expedition. (arial gletscher.) arion bjerg 70ø-376 (70°16.6´n 29°00.3´w). mountain 1010 m high in western gåseland, on the south side of snesø, so named during lauge koch’s 1958 expedition by eduard wenk’s greek assistant (j. papageorgakis), who was the only man to climb it. it was named after the greek singer arion, the semi-legendary poet and musician of mrthymna on lesbos, credited with the invention of the dithyramb (a choral poem or chant). arken 70ø-395 (70°28.8´n 29°43.5´w). nunatak in the glacier north of paul stern land. named by the 1963 geodætisk institut expedition for its supposed likeness to a ship in a sea of ice (arken = the ark). arkosedal – see øvre arkosedal. arkoseelv 73ø-293 74ø-203a (73°58.4´n 22°14.2´w). river in eastern hudson land draining ne to loch fyne. the name was used by gunnar säve-söderbergh during the 1931–34 treårsekspedi tionen in the form arkose river, and records the presence of arkosic sandstones. arkosepas 71ø-305 (71°34.0´n 24°48.2´w). pass between øvre arkosedal and ødemarksdal, on the west side of schuchert flod. named by enrico kempter during lauge koch’s 1956–58 expeditions. army iskappe 76ø-318 (76°52.0´n 24°14.3´w; map 4). ice cap in central dronning louise land. named by the 1952–54 british north greenland expedition to commemorate the help given by the british army, which provided several of the expedition members, as well as tracked vehicles and clothing. arnljotstua 75ø (75°08.3´n 20°30.4´w). norwegian hunting hut built in september 1932 for sigurd tolløfsen’s expedition on the south side of kap buch, easternmost c.h. ostenfeld land. it was named after arnljot tolløfsen, who died during a hunting trip in may 1933. arnold escher land 73ø-420 (74°00.0´n 28°15.0´w; maps 2, 4). nunatak group nw of andrée land. named during lauge koch’s 1951 expedition by hans r. katz for the pioneer of swiss geology, arnold escher von der linth [1807–72]. a structural geologist and stratigrapher, noted for his studies in the voralberg, he was professor of geology at zurich from 1852. the name was originally used in the incorrect form alfred escher land, which also appeared on the geodætisk institut 1:250 000 scale topographic maps. âronip sarpâ – see aaronip sarpaa. arringgletscher 72ø-350 (72°10.1´n 22°20.2´w). small glacier on se traill ø. named by h.p. heres during lauge koch’s 1956–58 expeditions. arthur dal 73ø-630 (73°03.6´n 26°57.9´w; map 4). glacier-filled valley in northern suess land, named during the 1931–34 treårsekspeditionen by eugène wegmann as arthur valley. arundel gate 72ø (72°07.5´n 24°33.8´w; map 5). mountain on the east side of bersærkerbræ, nw of panoramic peak, stauning alper. climbed by the 1982 university of sheffield expedition. arundel ø 73ø-12 (73°45.9´n 20°04.4´w; map 4). small island off the coast of hold with hope, named by william scoresby jr. in 1822 as cape arundel in compliment to the revd john arundel, who had married a sister of scoresby’s first wife (mary eliza lockwood). scoresby’s cape was probably a mountain on hold with hope, and the name was transferred to an island by white (1927). (arundel island). arundelhytten 73ø (73°46.0´n 20°04.9´w). danish hunting hut on arundel ø, off the coast of hold with hope, built by nanok in august 1949. arve 73ø-306 (73°42.0´n 22°26.4´w; map 4). river in eastern hudson land draining from afgrunden into storelv. named by heinrich bütler during lauge koch’s 1936–38 expeditions for the river of the same name in the mont blanc area of the french alps. arvehytten 73ø (73°41.6´n 22°09.6´w). danish hunting hut in stordal, hudson land, where the river arve flows into storelv. built by nanok in may 1947, it has also been known as vuachehytten and storelvhytten. arwidsson ø 72ø-28 (72°23.7´n 25°13.2´w; map 5). island at the confluence of alpefjord and forsblad fjord. it was named during a.g. nathorst’s 1899 expedition after ivar arwidsson [1873– 1936], the expedition zoologist, who subsequently became conservator at the zoological museum in uppsala. (arwidssons ö, arwids sonöya). aschenbrennerfjellet 72ø (72°56.2´n 23°50.3´w). mountain 1370 m high on western geographical society ø. the name is used on the nsiu maps of lacmann (1937), and was named after claus aschenbrenner [b. 1894], a german engineer who constructed photogrammetric instruments in munich and berlin. he also took part in the arctic flight of the ‘graf zeppelin’ in 1931. askers-øyane 72ø (72°12.7´n 23°40.4´w). name occasionally used in diaries of the 1930–32 møre greenland expedition (rogne 1981) for the present archer øer, on the south side of kong oscar fjord. askheimfjellet 72ø (72°57.6´n 24°15.1´w). mountain 1600 m high on western geographical society ø. used only on nsiu maps (lacmann 1937), it was named after thor askheim [b. 1889], a norwegian surveyor who took part in nsiu expeditions to east greenland 1931–33. assutsund [agsutsund] 77ø-69 (73°32.0´n 20°08.0´w; map 4). sound sw of c. silverberg ø in the inner part of skærfjorden. named by david malmquist during the 1931–34 treårsekspedi tionen for the motorboat asut, which was wrecked north of holland ø in october 1933. the name means ‘fast’. astartedal 70ø-50 (70°42.3´n 25°17.6´w). valley on the east coast 126 of milne land between charcot havn and kap leslie. named astarte-tal during the 1931–34 treårsekspeditionen by hermann aldinger, for the fossil lamellibranchs. (astarte valley). astarteelv 70ø-140 (70°36.9´n 22°39.9´w). river in astartekløft on the west side of hurry inlet. the name was first used in the form astarte river by harris (1931), reporting work during lauge koch’s 1926–27 expedition. the name was given for the abundant fossils. astartekløft 70ø-139 (70°36.9´n 22°39.9´w). ravine on the west side of hurry inlet in which astarteelv flows. the name derives from work by tom harris and alfred rosenkrantz during lauge koch’s 1926–27 expedition, and was commonly used in the form astarte klöft. (astarteklöft). astralhytten 75ø (75°49.9´n 19°39.7´w). norwegian hunting hut on the south side of sønderelv, on the coast of hochstetter forland about 12 km north of haystack. it was built by arktisk nærings drift about 1948–49 as a replacement for the 1933 sønderelv hut. astrupfjellet 73ø (73°59.3´n 22°25.5´w). mountain ridge 1700 m high in the nørlund alper, northern hudson land. the name is used on the nsiu maps of lacmann (1937), and commemorates eivind astrup [1871–95], a norwegian explorer who took part in robert peary’s 1891–92 and 1893–94 expeditions to north green land. astrup died mysteriously while skiing alone in norway in late december 1895. asut havn 74ø (74°59´n 21°44´w). sheltered bay east of mågenæs, on the north side of central grandjean fjord. the name was used in reports by helge g. backlund on his work during the 1931–34 treårsekspeditionen (in: koch 1955), and was given for the motorboat asut which used the bay as an anchorage in august 1932 (asut = swift). see also assutsund. atanikertik 70ø-222 (70°39.9´n 21°24.8´w). low ridge connecting the main part of rathbone ø, off se liverpool land, with its eastern headland. the name was recorded by the 1955 geodætisk institut name registration, and translates as ‘it has a little appen dage’. ataûsek âjertok 70ø (70°30.5´n 22°36.5´w). name used by rosenkrantz (1934) for a small hill or peninsula immediately south of tancrediakløft, on the west side of hurry inlet. it translates as the ‘bad place on the coast’. (ataûseq âjertoq.) attestupan – see ättestupan (near the end of this catalogue; ä is treated as æ in danish, and comes after z in the danish alphabet). attilaborgen 72ø-320 (72°00.2´n 25°19.5´w; map 5). mountain between sefström gletscher and krabbegletscher, northern stau ning alper. named by john haller and malcolm slesser following explorations during lauge koch’s 1954 expedition and slesser’s 1958 expedition. the slesser expedition climbed to within 30 m of the summit, and the first complete ascent was by the 1963 cambridge university expedition. it is a ferocious-looking mountain resembling a castle, appropriately named after attila, king of the huns from 434–453, and one of the greatest of barbarian rulers. (attilaborg.) aucellabjerg 74ø-154 (74°31.1´n 20°24.3´w; map 4). mountain 985 m high in western wollaston forland, named during the 1931–34 treårsekspeditionen by hans frebold for finds of species of aucella, a fossil lamellibranch. aucellabjerget 76ø-157 (76°07.0´n 18°38.1´w; map 4). mountain on southern store koldewey. so named by the 1906–08 danmarkekspeditionen for the presence of conglomerate beds containing the fossil aucella. it was occasionally called muslingebjerget (j. løve, personal communication 2009). (aucella mt.) aucellaelv 70ø-99 (70°37.4´n 23°30.5´w; map 4). river in sw jameson land. named by g.c. amdrup’s 1898–1900 expedition, the name appeared first in the form aucella elv in a report by nordenskjöld (1907). it was given for the fossil finds. (aucella river, aucellariver.) aucellaelv 74ø (74°30.6´w 20°29.3´w). river draining the slopes of aucellabjerg, wollaston forland. the name has been used as a reference locality by scientists visiting zackenberg forsknings station (e.g. meltofte & thing 1996). aucellapasset 74ø-93 (74°32.6´n 20°27.5´w; map 4). pass between palnatoke bjerg and aucellabjerg in wollaston forland, so named by lauge koch’s 1926–27 expeditions for the common fossil au cella. (aucellapas.) aucellaskråningerne 74ø (74°30.1´n 20°29.5´w). sw slope of aucel la bjerg. the name has been used as a reference locality by scientists at zackenberg forskningsstation. augpaleqisâp kûa – see aapaleqisaap kuaa. augpaleqisâq, augpaleqisâq kiáteq – see aappaleqisaaq, aappa leqi saaq kiatteq. augsburger spids 71ø (71°54.8´n 25°18.2´w; map 5). mountain on the north side of uppermost duart gletscher, central stauning alper. first climbed by karl m. herligkoffer’s 1966 expedition on 17 august, and named after the bavarian city of augsburg. (augs burger-spids.) augustadalen – see dronning augustadalen. augustadalshytten 74ø (74°24.2´n 19°09.5´w). name often used for the norwegian hunting hut built at the mouth of dronning augustadalen in july 1928 by the hird expedition. it was originally known as bjørnebu, and has also been called stordalen. (augusta dal hytten, dronning augusta dal hytten.) aumarssuit, aumarssuit nûat – see aamarsuit, aamarsuit nuaat. auspiciedalen 74ø-303 (74°06.7´n 21°00.5´w). small valley on south clavering ø, east of eskimonæs. the name originated from the wintering party at eskimonæs during the 1931–34 treårs ekspe di tionen, and refers to an area thought to be promising (auspicious) for mineral prospecting. o. eklund and david malm quist opened a small mine in the valley in 1933 in a pyrite vein with a distinct yellow gossan traceable for 1500 m. it contained 90 per cent pyrite and trace amounts of gold and silver (eklund 1944; harpøth et al. 1986). avalanche valley 72ø (c. 72°00´n 23°06´w). name used by ingstad (1937) for a valley visible from antarctic havn, eastern scoresby land. exact location uncertain. avantpost 74ø (74 °17.5¢n 20°39.8¢w). mountain ridge on eastern clavering ø. the name is found on lacmann’s (1937) maps. aztekerborgen 72ø-260 (72°21.3´n 24°38.5´w; map 5). mountain in the northern stauning alper, south of the front of skjold ungebræ. named by john haller during lauge koch’s 1954 expedition, for a resemblance to an aztec pyramid. it was first climbed by a norwegian party in 1951. aagenoesfjellet 74ø (74°22.1´n 20°46.7´w). mountain on north clavering ø, so named on lacmann’s (1937) topographic maps. b bacchus gletscher 71ø (71°44.0´n 25°38.1´w; map 5). glacier in the ne part of the borgbjerg gletscher region, southern stauning alper, north of bacchustinde. probably named by the 1977 schwä bische stauning alper expedition. bacchustinde 71ø (71°43.3´n 25°39.4´w; map 5). peak 2397 m high in the ne part of the borgbjerg gletscher region, southern stauning alper. probably first climbed and named by the 1977 schwäbische stauning alper expedition. bach dal 76ø-319 (76°52´n 23°33´w; map 4). valley in central dronning louise land draining sw from ad astra iskappe to join beethoven dal. one of the names given by the 1952–54 british north greenland expedition for german composers, it commemorates johan sebastian bach [1685–1750], noted especially for his choral music. backlund bjerg 71ø-342 72ø-134 (71°57.7´n 28°11.0´w). moun tain c. 1600 m high on the north side of innermost nord vestfjord. named by eduard wenk after helge götrik backlund [1878– 1958], a swedish geologist especially noted for studies of rapakivi 127 granites. he was professor of geology at uppsala from 1924–43. back lund was said to have been the first to set foot on the mountain in august 1934, and was one of a party that included wenk and narrowly escaped disaster when giant waves produced by the calving of daugaard-jensen gletscher swamped their boat. wenk climbed to the highest point of the mountain in 1954. the mountain has also been called reinhard bjerg. backlund ridge 72ø 73ø (73°00.3´n 23°06.9´w). mountain ridge on northern geographical society ø, east of rudbeck bjerg. the name was used by gunnar säve-söderbergh (1933, plate 3) during his work on the 1931–34 treårsekspeditionen, and commemorates helge g. backlund. see also backlund bjerg. (backlundkammen.) badger 71ø (71°08.7´n 26°46.1´w). summit 2044 m high on the ice cap between catalinadal and edward bailey gletscher, renland. climbed and named by the 2007 west lancashire mountaineering group expedition. badlanddal 73ø-41 (73°34.0´n 21°48.0´w; map 4). broad n–strending valley between loch fyne and mackenzie bugt. so named by lauge koch’s 1929–30 expeditions in the form badland valley, because of the characteristic erosion forms developed in the glacial sediments on the valley floor. (badland tal.) baesdalen 74ø-245 (74°09.5´n 20°36.3´w). valley on se clavering ø, between rundetårn and brinkley bjerg, in which moskus okseelv flows. the name originated from the wintering party at eskimonæs during the 1931–34 treårsekspeditionen, and was first used in the form basdalen (baes = bas = lead dog in a sledge team). mausdalen has also been used. bagatellerne 79ø-6 (79°39.9´n 18°02.0´w). group of small islands off se hovgaard ø. so named by the 1906–08 danmark-ekspe ditionen, which left a depot here in october 1907. named for their small size (bagatel = trifle). bagdalen 80ø-53 81ø-67a (80°48.0´n 17°07.5´w; map 4). name given to a valley that appeared to run west and north of amdrup land (bag = behind), so named by eigil nielsen during the 1938–39 mørkefjord expedition. to the north the valley follows a broad depression in the ice cap. to the west the valley drains into ingolf fjord on the east side of tobias gletscher. bagfjorden 76ø-117 (76°34.6´n 22°22.5´w; map 4). name given by j.p. koch’s 1912–13 expedition to a small fjord unexpectedly found west of lindhard ø and kap jarner (bag = behind), partially blocked by an ice tongue from storstrømmen. (bagfjord, bak fjörður.) baie brongniart 69ø (69°14.5´n 25°06.0´w). bay on the northern blosseville kyst, probably identical with barclay bugt. the name is found on a map by jules de blosseville from 1833 (fig. 4). baie watkins 70ø (70°36.5´n 21°46.1´w). name used for a small bay on the west side of lillefjord, southern liverpool land, on a very inaccurate map prepared by maurice parat during j.b. charcot’s 1933 expedition (parat & drach 1934). an adjacent glacier was referred to as glacier watkins. the names commemorate henry george (gino) watkins [1907–32]. see also watkins bjerge. baie d'orleans 77ø (c. 77°30´n 19°30´w). large bay bounded to the east by kap philippe, corresponding approximately to the present skærfjorden. the name was given during the 1905 duke of orléans expedition. see also hertugen af orléans land. bakkehaug 74ø (74°25.8´n 21°26.9´w). norwegian hunting hut on the nw side of clavering ø, west of louise elv, built by the foldvik expedition in august 1927. it has also been called louise elv hytten. (bakkehytten, bakkehuset.) bakkehuset – see karlsbak. bakkehytta 72ø (72°59´n 24°32´w). norwegian hunting hut built in september 1930 for arktisk næringsdrift on the nw point of geographical society ø. the name derives from it being placed on an awkward slope (= bakke). the hut has also been known under the names svedenborg, joplassen, røvballehytten and valborghytten. balderbreen 74ø (74°15.0´n 21°02.3´w). glacier on central clav ering ø, a minor tributary to skillegletscher. so named on the nsiu maps of lacmann (1937) after balder, son of odin in old nordic mythology, noted for his gentleness. balders hage 74ø (74°59´n 21°45´w). name reported used by the wintering party at kulhus in 1935 for mågenæs, the peninsula on the north side of central grandjean fjord. see also balderbreen. baldwin-huset 74ø (74°56´n 17°37´w). name sometimes used for the eight-sided depot hut built at kap philip broke in south shannon for the 1901 baldwin-ziegler expedition. it has also been known as ziegla-husa. see also kap philip broke. balmunggletscher 74ø-381 (74°24.1´n 21°05.4´w). small glacier on northern clavering ø, named on the nsiu maps of lacmann (1937) in the form balmungbreen. the name is derived from old nordic mythology. balnes 74ø (74°20.1´n 21°56.2´w). small cape on the east coast of payer land, south of revet. named after the area near troms, norway, from which many norwegian hunters originated. lac mann’s (1937) maps use the spelling balsnes. baltos bre 71ø (71°54.5´n 25°11.7´w; map 5). name given to a northern branch of roslin gletscher by the 1996 norwegian stauning alper expedition. it was named after samuel johannes balto [1861–1921], who accompanied fridtjof nansen on his crossing of the inland ice in 1888. balås flyveplass 73ø (c. 73°27´n 21°48´w). natural landing field on vestersletten, west of mackenzie bugt. named during the 1932 nsiu expedition, which had two aircraft used mainly for aerial photography based there. (balås flyvplass.) bamsegletscher 72ø-157 (72°17.6´n 22°35.4´w). glacier on se traill ø, south of mountnorris fjord. named during lauge koch’s 1936–38 expeditions by hans peter schaub for the polar bear (bamse = teddy bear). banangletscher 72ø-342 (72°18´n 22°54´w). glacier on se traill ø. named by hans-peter heres during lauge koch’s 1956–58 expeditions for the banana-like shape of the glacier. barclay bugt 69ø-12 (69°14.5´n 25°06.0´w; map 3). name given by amdrup’s 1898–1900 expedition to the bay on the northern blosseville kyst south of kap barclay. it may be identical with jules de blosseville’s 1833 baie brongniart. (barclay-bugt, barclay bight, barchlay bugt.) bärenkamm 73ø (73°22.5´n 26°05.0´w). mountain in southern andrée land. the name is found on a panorama sketch drawn by john haller in 1949 published in schwarzenbach (1993). barenzahne 72ø (72°04.1´n 25°13.3´w; map 5). mountain between sefström gletscher and gully gletscher, stauning alper. climbed and so named by the 1964 zurich expedition. barnabas dal 73ø-639 (73°05.9´n 23°56.7´w). valley on southern ymer ø, draining se to sofia sund. named during the 1931–34 treårsekspeditionen by ove simonsen after jørgen barnabas, a greenlander who assisted the expedition from 1932 to 34 and sometimes hunted here. it has also been called raudalen. barnabasdal hytten 73ø (73°04.1´n 23°43.3´w). norwegian hunting hut built in october 1930 on the east side of barnabas dal, ymer ø. it is also known as raudalshytta and stor-dalen. barnacle cliffs 77ø (77°36.6´n 20°48.7´w). cliff west of klægbugt, nordmarken, where barnacle geese nest. named by the 1987 irish expedition to northern east greenland. barriere gletscher 74ø-375 (74°43.3´n 21°59.7´w). minor glacier in western svejstrup dal, the valley between th. thomsen land and a.p. olsen land, so named by the 1948 leeds university expedition because it formed a difficult obstacle. (barrier glacier). barrieren 73ø-428 (73°20.2´n 24°47.9´w). elongate n–s mountain ridge on northern ymer ø, named by silvio eha during lauge koch’s 1947–49 expeditions (barrieren = the barrier). barrieren 76ø-336 (76°23.4´n 25°54.8´w; map 4). high peak in dronning louise land, which to the 1952–54 british north greenland expedition appeared to form a barrier across budolfi isstrøm when sledging down that glacier from the west. 128 barrikadegletscher 72ø-151 (72°19.0´n 22°53.0´w). glacier on se traill ø, ne of steenstrup bjerg. so named during lauge koch’s 1936–38 expeditions by hans peter schaub because it formed a barrier across bjørnedal in 1937; the glacier has retreated significantly since then. barrikadental 72ø (72°20.9´n 22°47.3´w). name used by stauber (1938) for the present bjørnedal, a valley on the south side of mountnorris fjord on se traill ø partially blocked by barrikade gletscher. barth-hytta 75ø (75°24.8´n 21°11.3´w). norwegian hunting hut built in august 1932 for john giæver’s expedition on the north side of ardencaple fjord, west of the barth bjerge. it was originally called berglann, and more usually known as holmsnes. now a ruin. barth bjerge 75ø-18 (75°29.0´n 20°44.0´w; map 4; see also fig. 59). mountain range west of hochstetter forland, on the north side of ardencaple fjord. named by karl koldewey’s 1869–70 expedition as barth berge, probably after heinrich barth [1821–65], an influential german geographer who made important expeditions to central africa. (barth mountains, barth bjergene, barthberge, barth berge, barthfjellene.) barth bjerge 75ø (c. 75°24´n 20°20´w). name used for the norwe gian hunting hut built in 1949 by arktisk næringsdrift in a valley on the north side of the barth bjerge, southern dronning mar grethe ii land. it was reported to have been destroyed by strong winds in 1949 or 1950 (p.s. mikkelsen 1994, 2008). bartholin borg 74ø-394 (74°22.2´n 24°21.6´w; map 4). mountain c. 1600 m high in east bartholin land. named by john haller during lauge koch’s 1956–58 expeditions. thomas bartholin [1616–80] was a danish mineralogist noted for studies of amber in denmark (borg = castle). bartholin bræ 69ø-19 (69°38.0´n 24°04.0´w). glacier west of henry land, on the northern blosseville kyst. named by g.c. amdrup’s 1898–1900 expedition as bartholins bræ for thomas bartholin. see also bartholin borg. böggild (1905) used henry glacier for the same feature. (bartholin glacier.) bartholin land 74ø-139 (74°24.5´n 25°00.0´w; maps 2, 4). land area on the north side of waltershausen gletscher. so named by lauge koch’s 1929–30 expedition, and better defined in 1932 as a result of aerial reconnaissance (fig. 15). its boundaries are now vibeke gletscher, indelukket, and to the west longitude 25°40´w. see also bartholin borg. (bartholins land, bartholin-hochland.) bartholin nunatak 74ø-391 (74°16.6´n 25°21.7´w; map 4). nuna tak south of korsgletscher, in southern bartholin land. named during lauge koch’s 1956–58 expeditions by john haller. see also bartholin borg. bartletts skær 74ø (c. 74°04´n 21°45´w). submerged skerry west of kap stosch where robert bartlett ran aground in the effie m. morrissey in 1931. he was dragged off with the assistance of the polarbjørn, and was somewhat offended to receive later a bill from the owners (see bartlett 1934). the name is used in den grønlandske lods (1968). robert a. bartlett [1875–1946] was one of the great american arctic skippers. he made more than 40 arctic voyages, and is particularly noted for his association with robert peary as first mate on the windward and skipper of the roosevelt (1898–1908). he was also skipper of the ill-fated karluk (1913–14), and made numerous scientific voyages to the arctic with the effie m. morrissey (1925–45). basalt havn 74ø (74°20´n 20°26´w). reference locality used by dunbar (1955) for the harbour (= havn) on the nw coast of clavering ø, west of basalt ø. dunbar misspells the name as basalt haven. (diabase haven.) basalt table mountain 70ø (c. 70°33´n 23°18´w). locality in southern jameson land, referred to by spath (1935) as a place where h. bütler collected fossil ammonites. basalt sills form small plateaus at several locations in this area (see aldinger 1935, plate 2). basaltdal 72ø-347 (72°14´n 22°37´w). valley on se traill ø, so named during lauge koch’s 1956–58 expeditions by hans-peter heres for the presence of thick basalt sills. basaltelv 70ø-403 (70°27.2´n 22°43.2´w). river in se jameson land, originally named by aldinger (1935) as basalt river for the dolerite sills. the name was officially approved in 1972, following new work in the area by ggu. basaltkap 74ø-164 (74°08.7´n 20°28.9´w; map 4). minor cape on the se coast of clavering ø. the name arose during the 1931–34 treårsekspeditionen, and was first used in a report by backlund & malmquist (1932). the cape is formed by a basalt intrusion. basaltkløft 72ø-380 (72°02.1´n 23°28.3´w). minor valley or ravine in northern scoresby land, west of the pictet bjerge. so named by hans kapp during lauge koch’s 1957–58 expeditions because of the presence of a basalt sill. basaltnæs [aamarsuit nuaat] 70ø-293 (70°27.4´n 22°16.1´w). peninsula east of kap hope, southern liverpool land, made of basaltic rocks. so named by alfred rosenkrantz during lauge koch’s 1926–27 expedition. basaltpynt 72ø-164 (72°31.9´n 22°11.1´w; map 4). cape on eastern traill ø, nw of the mouth of æbeltoft vig. named during lauge koch’s 1936–38 expeditions by hans peter schaub for the basaltic rocks. basaltpynten 74ø (74°20.3´n 20°26.4´w). peninsula on eastern clav ering ø, west of basaltø. the name is used on lacmann’s (1937) maps. basaltspids 74ø-387 (74°03.8´n 28°21.1´w). peak of basalt in northern arnold escher land, named by hans r. katz during his traverse through the nunatak region on lauge koch’s 1951 expedition. basaltsø 72ø-400 (72°43.3´n 22°29.3´w; map 4). lake on southern geographical society ø. the name came into use in the 1950s during lauge koch’s geological expeditions, and records the occurrence of basaltic rocks. hofgaardvatna has also been used. basaltø 74ø-116 (74°20.1´n 20°22.9´w; map 4). island in young sund off eastern clavering ø, so named during lauge koch’s 1929–30 expeditions in the form basalt island because it is composed of basaltic rocks. klippeø has also been used. baselbjerget 74ø-339 (74°51.1´n 20°23.8´w; map 4). mountain 750 m high on western kuhn ø, so named during lauge koch’s 1936–38 expeditions by wolf maync and andreas vischer for the swiss city of basel. (baselbjærget.) baselfjeld 73ø-705 (73°15.3´n 28°42.5´w). nunatak 2600 m high in western frænkel land. named by john haller and eduard wenk during lauge koch’s 1951 expedition, for the city and university of basel in switzerland where both were based. basen 74ø (74°27.9´n 20°38.4´w). name used for the house and depot hut built in 1947 at zackenberg bugt, east of zackenberg, for eigil knuth’s 1947–50 danish peary land expeditions; it was subsequently used by the 1952–54 british north greenland expedition. it is also known as zackenberg base. basisdalen 71ø-104 (71°36.8´n 22°14.5´w). common name for søndre basisdal and nordre basisdal in se canning land. the name appears to have first been used by säve-söderbergh (1937) in the form basis valley, and derives from work during lauge koch’s 1936–38 expeditions. basiskæret 76ø-104 (76°46´n 18°39´w). swampy area north and east of the original expedition house at danmarkshavn. so named by the 1906–08 danmark-ekspeditionen because the triangulation base for the maps of the expedition was measured here. the staff at danmarkshavn weather station know it as kærene (kær = marsh). (basiskær.) bass rock 74ø-18 (74°43´n 18°16´w; maps 2, 4). small island ne of lille pendulum. named by douglas clavering in 1823 for its resemblance to bass rock on the south side of the firth of forth, scotland, an impressive, steep-sided island which was the site of a castle, later a prison and fortress destroyed in 1694. depot huts were built on bass rock in 1901 (see bass rock-husene). the norwegian floren expedition climbed to the summit in june 1909. (bass klippe.) bass rock-husene 74ø (74°42.8´n 18°15.2´w). two eight-sided de pot huts were built on the south side of bass rock for the baldwinziegler expedition in 1901. they were subsequently visited and used by the 1906–08 danmark-ekspeditionen, the 1909–12 ala bama expedition, østgrønlandske fangstkompagni 1920–24, and nanok 1929–30. the alabama expedition made use of the supplies in the depot after the alabama sank in winter quarters off shannon, as did the crew of the dagny in 1921 after their ship had been crushed in the ice. the huts were transferred to norwegian ownership in 1930, and in 1969 to danish ownership when all norwegian huts in east greenland were taken over by denmark. they have also been referred to as the ziegler-husa. bastian bugt 74ø-25 (74°55.2´n 20°08.5´w; map 4). pronounced bay on eastern kuhn ø. named by karl koldewey’s 1869–70 expedition as bastians bai for adolf bastian [1826–1905], a german explorer and ethnologist who had formed a committee for raising funds for the expedition. a norwegian hunting hut (bolettestua) was built on the north side of the bay by the 1932–34 tolløfsen expedition. (bay of bastian, bastians bugt, bastiansbucht.) bastians dal 74ø (74°53.6´n 20°11.9´w). name occasionally used for the e–w valley on kuhn ø draining into the head of bastian bugt (e.g. vischer in: koch 1955). bastille 71ø (71°42.0´n 25°04.2´w; map 5). peak 1870 m high south of concordia, on the sw side of bjørnbo gletscher, southern stauning alper. first climbed by james clarkson’s 1961 expedition, and named after the medieval fortress on the east side of paris, a notorious french state prison in the 17th and 18th centuries. bastionbugt 72ø-122 (72°51.0´n 25°11.9´w). bay on nw ella ø, east of bastionen, so named during the 1931–34 treårsekspedi tionen by the ella ø wintering party. bastionen 72ø-50 (72°50.3´n 25°18.7´w; fig. 29). mountain forming the west cape of ella ø, which rises nearly vertically from the sea for 1200 m. named by a.g. nathorst in 1899 for its massive appearance. (bastion, mt. bastion). bastionerne 76ø-57 (76°55.6´n 20°08.5´w). small hills on the east side of lakseelven, western germania land. so named by the 1906–08 danmark-ekspeditionen. (the bastions, bastion, basti onen.) bastionerne 76ø (c. 76°36´n 18°48´w). name used for part of the east side of northern store koldewey during the 1906–08 dan mark-ekspeditionen by thostrup (2007). perhaps intended as descriptive rather than a place name (j. løve, personal communication 2009). bastionpynt 72ø-277 (72°50.6´n 25°21.0´w). cape on the west side of bastionen, which is also the westernmost point of ella ø. named by john cowie during lauge koch’s 1949–54 expeditions. bath elv 72ø-235 (72°27.1´n 22°27.1´w). river on eastern traill ø, draining south into mountnorris fjord. so named by desmond donovan during lauge koch’s 1949–50 expeditions for his home town of bath in england. bathosbjerg 73ø (73°32.3´n 25°44.9´w). mountain 2032 m high on the south side of grejsdalen, andrée land. climbed by the 2007 army boreal zenith expedition. bavariaspitze 72ø (72°01.0´n 24°58.0´w; map 5). mountain 2180 m high east of sefström gletscher, stauning alper. first climbed by hans gsellman’s 1957 expedition, and named as a friendly gesture to the german member of the party, herman köllensberger. (bavariaspids.) bavnen 74ø-291 (74°47.9´n 21°32.1´w; map 4). mountain 1250 m high between odin dal and svejstrup dal, th. thomsen land. the name originated from the wintering party at kulhus during the 1931–34 treårsekspeditionen (bavnen = the beacon). bay fjelde 70ø-56 (70°40.5´n 25°45.1´w; map 4). group of peaks up to 830 m high west of kap leslie, east milne land. named by alfred rosenkrantz during lauge koch’s 1926–27 expeditions as mts bays fjælde after edvard bay, geologist of the 1891–92 den østgrønlandske expedition led by carl ryder. see also edvard bay dal. bayerndom 72ø (72°08.5´n 25°42.4´w). mountain 2312 m high in the trekantgletscher area, west of alpefjord. it was climbed by wolfgang weinzierl’s 1970 expedition, and named after the south german district of bayern (bavaria). exact location a little uncertain. (bavarian cathedral.) bays elv 70ø (70°39.4´n 25°37.1´w). minor river draining the flanks of bay fjelde, se milne land, a tributary to aldinger elv. the name appears on the maps of callomon & birkelund (1980). attempts to obtain official approval of the name in 1977 were unsuccessful. see also edvard bay dal. bear mountain 71ø (c. 71°25´n 23°15´w). name used by ingstad (1937) for one of the summits se of olympen on jameson land where they shot a bear. exact location uncertain. they were stormbound in their camp in the upper reaches of pingel dal for eight days in 1932, and survived on a diet of almost raw bear meat. bear peak 72ø (72°07.7´n 24°43.8´w) peak about 800 m north of tintagel fjeld on the west side of bersærkerbræ, northern stauning alper. a rock perched on the summit resembles a bear in shape. climbed and so named by the 1991 scottish stauning alper expedition. (bear.) fig. 29. looking south at bastionen, the cliff forming the west point of ella ø. kap alfred is the north cape of lyell land, separated from ella ø by narhvalsund. kongeborgen in western traill ø can be seen in the background. the john haller photograph collection, geus archive. traill ø kongeborgen bastionen ella ø narhvalsund kap alfred lyell land 129 130 beaufort tinde 72ø-361 (72°01.5´n 25°09.2´w; map 5). rock spire 2277 m high in the northern stauning alper, ne of sefström gletscher. first climbed by malmolm slesser’s 1958 expedition, and named after beaufort castle, invernesshire, a 19th century man sion, seat of the frasers of lovat. hans gsellman’s 1957 expedition had earlier reached to within 100 m of the summit, and called it kapellenturm. (beaufort.) beaumaris fjeld 72ø-491 (72°06.7´n 24°36.0´w; map 5). moun tain 1900 m high at the head of bersærkerbræ, northern stauning alper. first climbed by john hunt’s 1960 expedition, and named after beaumaris castle, anglesey, north wales. the second ascent was by the 1968 queen mary college expedition. the position of this mountain is incorrect in bennet’s (1972) guide to the stauning alper, and has caused problems for many climbing groups. some later climbers viewed the higher peak to the east as the possible ‘real’ beaumaris fjeld, and beaumaris fjeld was then labled incorrectly as beaumaris west. (beaumaris.) beaumaris gletscher 72ø-492 (72°07.2´n 24°37.0´w; map 5). glacier on the south side of bersærkerbræ, north of beaumaris fjeld, northern stauning alper. named by john hunt’s 1960 expedition. beaumaris west – see beaumaris fjeld. beethoven dal 76ø-320 (76°47.8´n 23°37.2´w; map 4; fig. 21). valley in central dronning louise land. one of the names given by the 1952–54 british north greenland expedition for german composers, it was named after ludvig van beethoven [1770–1827], noted especially for his classical symphonies. begtrup vig 72ø-82 (72°26.3´n 22°18.4´w). bay on the north side of mountnorris fjord, eastern traill ø. named during the 1931–34 treårsekspeditionen by ove simonsen for the danish locality of the same name in the mols district of jylland. beinhaugen 72ø (72°31.2´n 24°39.5´w). norwegian hunting hut at kap lagerberg, se lyell land, built by the møre expedition in august 1930. the name (= bone hill) is a reference to inuit remains near the hut. it is now generally known as lagerberghytte. (bein hauen.) belgica banke 78ø-41 (c. 78°09´n 18°00´w; fig. 30). offshore bank discovered during the 1905 duke of orléans expedition, and named bank de la belgica for the expedition ship the belgica, a 300-ton three-masted barque. (belgica shoal.) bellavista 73ø-363 (73°45.5´n 25°12.5´w). mountain in western strindberg land, overlooking geologfjord, named by hans r. katz during lauge koch’s 1948–49 expeditions for the spectacular view. bellerophon glacier 71ø (71°49.1´n 25°21.0´w). minor tributary glacier to bjørnbo gletscher, southern stauning alper. named by james clarkson’s 1961 expedition for bellerophon, the hero of the iliad. bellevue 71ø-265 (71°58.1´n 24°06.7´w; map 5). mountain in the werner bjerge between langefirn and bredefirn. the name ap pears to have been given by the place name committee as a replacement for styger’s (1951) pyramiden. the mountain was climbed by hans stauber in 1948 and peter bearth in 1953. ‘bellevue’ is a common locality name in switzerland. bendaelv 73ø-192 (73°37.6´n 21°48.4´w). river flowing into the south end of loch fyne. so named on the nsiu 1932 map (nsiu 1932a), as benda because of the pronounced curve in the inland course of the river. benjamin dal 73ø-640 (73°20.9´n 25°42.2´w). valley in se andrée land, draining into eleonore bugt west of teufelsschloss. named by ove simonsen during the 1931–34 treårsekspeditionen for benjamin samuelsen, a greenlander who assisted the surveying parties. (benjamins dal.) benjamins bugt 73ø (73°23.9´n 25°30.6´w). name used by the 1972 university of dundee expedition for the bay at the mouth of benjamin dal, which is part of eleonore bugt. bennethytta 73ø (73°22.6´n 21°41.8´w). norwegian hunting hut on the south side of kap bennet, eastern gauss halvø, built by the foldvik expedition in august 1927. it has also been known as giesecké, giskehytta and foldvik. (kap bennet hytte.) bennethøgda 73ø (73°24.9´n 21°40.7´w). name used on the nsiu (1932a) map for the 358 m hill west of kap bennet, eastern gauss halvø. (bennet ridge.) berchtesgadener gletscher 71ø (71°54.8´n 25°36.1´w). name used by the 1967 berchtesgadener expedition for the glacier on the west side of spærregletscher, stauning alper, more usually known as hecate gletscher. named with berchtesgadener kopf at the head of the glacier for berchtesgaden, a popular holiday and climbing resort in the bavarian alps, germany. berchtesgadener kopf 71ø (71°52.6´n 25°40.0´w; map 5). mountain fig. 30. the steam-assisted sailing ship belgica used by the baldwin-ziegler expedition to lay out depots in 1901 and by the duke of orléans in 1905. the belgica, formerly patria, was purchased by adrien de gerlache de gomery in 1896 for the 1897–99 belgian antarctic expedition. 131 about 2500 m high between prinsessegletscher and hecate gletscher, stauning alper. named and first climbed by the 1967 berchtes gadener expedition. berchtesgadener tinde 71ø (71°50.0´n 25°31.1´w; map 5). peak 2560 m high on the south side of the upper basin of spærre gletscher, stauning alper. climbed by karl m. herligkoffer’s 1966 expedition on 18 august, and named after the home town of josef anzenberger, one of the climbers. see also berchtesgadener gletscher. berg fjord 76ø-34 (76°34.0´n 18°55.5´w; map 4). fjord on the west side of store koldewey, which nearly divides the island into two parts. named by the 1906–08 danmark-ekspeditionen as bergs fjord, for the chairman of the engineers’ association in copen hagen (thostrup 2007), who had helped obtain permission for hermann koefoed’s participation in the expedition (j. løve, personal communication 2009). berg fjordhytten 76ø (76°35.1´n 18°49.5´w). norwegian hunting hut, built in september 1938 in the ne part of berg fjord, store koldewey, by the norsk–franske polarekspedisjon. it is also known as inderhytten. bergfjordhytten 76ø-202 (76°35.7´n 18°44.7´w). danish hunting hut on the east side of store koldewey, at the col leading to berg fjord; it is also known as pashytten and yderhytten. it was built by nanok in august 1933. (bergs fjord hytten.) berggeistspids 71ø (71°51.0´n 25°33.5´w; map 5). peak about 2615 m high on the sw side of the upper basin of spærregletscher, stauning alper. climbed by karl m. herligkoffer’s 1966 expedition, and named after their climbing club. berglann – see holmsnes. berlin-stua 74ø (74°40.0´n 19°19.7´w). norwegian hunting hut in the bay se of kap berlin, northern wollaston forland. built by the møre expedition in august 1930. (kap berlinhytte.) berliner bjerg 71ø (71°53.2´n 25°32.9´w; map 5). mountain on the west side of the upper basin of spærregletscher, stauning alper. first climbed by karl m. herligkoffer’s 1966 expedition on 23 august, and named after the city of berlin, germany. position uncertain; some climbers consider this peak may be identical with schneekuppel. bern plateau 74ø (74°36´n 19°30´w). name used by maync (1947, 1949) for the basalt-capped southern part of brorson halvø in northern wollaston forland. named during lauge koch’s 1936–38 expeditions for the city of bern, switzerland. bernbjerget 74ø-338 (74°47´n 20°21´w; map 4). mountain 620 m high on south kuhn ø, so named during lauge koch’s 1936–38 expeditions by wolf maync and andreas vischer, for the swiss city of bern (maync 1947). (bernbjærget.) bernhard studer land 74ø-384 (74°04´n 27°10´w; map 4). nunatak region between eyvind fjeld gletscher and hindrings gletscher, north of andrée land. named during lauge koch’s 1951 expedition by hans r. katz after bernhard rudolf studer [1794–1887], a noted swiss geologist. he was professor of geology at the university of bern from 1834, and noted for his stimulation of the first geological mapping of switzerland and studies of molasse. (bernhard studers land.) bersærker tinde 72ø-372 (72°04.4´n 24°46.1´w; map 5). domi nant peak 2428 m high at the head of bersærkerbræ, north stau ning alper. the name is attributed to john haller and malcolm slesser, and derives from the adjacent glacier. it was first climbed by the 1968 queen mary college expedition. (bersaerker tinde.) bersærkerbræ 72ø-98 (72°08.0´n 24°38.0´w; map 5; fig. 31). large glacier in the northern stauning alper draining ne into skeldal. named by ove simonsen during the 1931–34 treårsekspe ditionen. in old nordic mythology the ‘bersærker’ (‘bare’-breast) were savage warriors who in their frenzy in battle destroyed everything in their path. berzaerkerspire 72ø (72°07.7´n 24°47.3´w; map 5). dramatic 2000 m high peak, officially known as spiret, between dunottar gletscher and bersærkerbræ. this name is invariably used by climbers in preference to the official name. (berzaerker spire.) berzelius bjerg 72ø-36 (72°28.0´n 25°05.0´w; maps 4, 5; fig. 32). mountain massif in se lyell land. named by a.g. nathorst in 1899 as berzelius' berg, or berzelii berg. jöns jakob berzelius [1779–1848] was a noted swedish chemist, the father of modern chemistry, most celebrated for his table of atomic weights published in 1818. (berzelius mountain, berzelius peak, berzelius bjærg, berzelius fjellet). bessel fjord 75ø-6 76ø-211a (75°59´n 21°00´w; maps 2, 4). fjord between ad. s. jensen land and dronning margrethe ii land. it was named bessel bai by karl koldewey’s 1869–70 expedition after franz friedrich wilhelm bessel [1784–1846], a noted german astronomer who was professor of astronomy and director of the observatory at the university in königsberg. koldewey apparently applied the name to a bay at the mouth of the fjord, and it was moved to the fjord itself by the 1906–08 danmark-ekspeditionen. the bessels fjord form (with final genetive ‘s’) is most often encountered, and was that used on the geodætisk institut 1:250 000 scale topographic map sheets up to 1970. (besselbai, bessel bay, bessel fjord.) bessel fjord 76ø-211 (76°03.4´n 20°06.0´w). danish hunting station at the mouth of trumsdalen on the north side of the mouth of bessel fjord. built by nanok in 1932, it replaced a hut on the same site (bessels fjord hytten) built in may 1931. the station was only manned in 1932–33, as the site proved liable to constant strong winds. the station was still standing in 1989, but is now in poor condition. it has also been known as trumsdalen. (bessels fjordstation, bessel fjord station.) bessels fjord hytten – see bessel fjord. besselsfjordhytten 75ø (75°56.0´n 19°56.5´w). danish hunting hut built for nanok in september 1932 at kap møbius, south of the mouth of bessel fjord. it has the approved name mundingshytten. bessfjellet 73ø (73°22.2´n 22°14.5´w). mountain in the southern giesecke bjerge, gauss halvø, corresponding to the present huit feldt bjerg. so named on the nsiu 1932 map (nsiu 1932a), the name derives from a norwegian dialect form (besse = male bear). the mountain lies north of the björnedalen of the nsiu 1932 map (nsiu 1932a). bessvatnet 74ø (74°13.6´n 22°12.1´w). lake on the se flank of blosseville bjerg, at the front of wordie gletscher. used only on stauning alper bersæ rke rb ræ fig. 31. view southwards over bersærkerbræ to the high summits of the stauning alper. the john haller photograph collection, geus archive. 132 nsiu maps (lacmann 1937), the name derives from the norwe gian dialect word for a male bear (= besse). beta nunatak 71ø-381 (71°48.5´n 29°58.0´w; map 4). largest nunatak in the alfabet nunatakker, western charcot land. named during the 1967–72 ggu scoresby sund expeditions. see also alfabet nunatakker. betulahavn 75ø-59 (75°00.8´n 22°03.3´w). bay with a good anchorage on the sw side of central grandjean fjord. the locality was visited by gunnar seidenfaden in 1932 during the 1931–34 treårsekspeditionen, and samples of dwarf birch (betula nana) were collected. the name was used as a botanical reference locality (gelting 1934) and records the then northernmost occurrence of the species. (betula harbour.) betulahavnhytten 75ø (75°01.1´n 22°03.5´w). danish hunting hut at betulahavn, inner grandjean fjord, built by nanok in 1951. it is also known as birkedalshytten. (betula havn hytten.) betvatna 72ø (72°42.8´n 21°58.0´w). small lake on eastern geo graphical society ø, on the peninsula lacmann (1937) called weren skioldflya. the lake was named after elisabeth (beth) ma thilde werenskiold [b. 1897], wife of the painter dagfin weren skiold. see also dagfinvika. (bethvatna.) beurmann – see olestua. bielven 70ø (70°54.8´n 22°24.9´w). name used by g.c. amdrup’s 1898–1900 expedition for the tributary to ryder elv which drains hodal in liverpool land. big chocolate mountain – see chokoladebjerg. big nev 70ø (70°48.2´n 21°55.7´w). peak 761 m high in liverpool land, west of innermost horsens fjord. it was climbed and named by the 2002 loughborough grammar school expedition. big river 72ø (72°31.4´n 23°59.4´w). name used by the 1974 joint biological expedition for a river west of karupelv draining into holm bugt, sw traill ø. bildsøe nunatakker 77ø-99 78ø-18 (78°05.0´n 23°40.0´w; maps 1, 2, 4). nunatak group west of hertugen af orléans land, named by the 1909–12 alabama expedition as bildsöe’s nunatakker. jens arnold diedrick jensen bildsøe [1849–1936] was noted for five exploration voyages to west greenland, four of them as leader, that included a 70 km sledge expedition on frederikshåb isblink. bildsøe was navigation director at marstal navigation school when ejnar mikkelsen was studying there (j. løve, personal com munication 2009). binnenland 74ø 75ø (74°30´–75°30´n 20°20´w). name used in the minutes of the publication committee of karl koldewey’s 1869–70 expedition (verein für die deutsche nordpolarfahrt 1870–76). it was apparently used for kuhn ø and the adjacent mainland in the sense of the ‘interior’ of the country (binnen = within). binnental – see inderdal. binucleus 74ø-129 (74°12.7´n 20°49.6´w). mountain on se clavering ø with two summits, 1493 m and 1471 m high. named during lauge koch’s 1929–30 expeditions in the form mt. binu cleus. see also monacleus and trinucleus. (binucleusfjellet, binu cleus bjerg ). biot-stua 71ø (71°57.0´n 22°44.1´w). norwegian hunting hut 3 km nw of kap biot, eastern scoresby land, built by the møre expedition in august 1930. it also goes by the names davy sund hytten, villa and nordre biot. birgitnæs 74ø (74°08.7´n 20°28.9´w). minor cape on se clavering ø, possibly the present basaltkap. so named on a sketch map in gustav thostrup’s 1921 logbook (møller 1939). girl’s name. birgitsbjærg 72ø (72°20.5´n 24°33.1´w). name given by erdhart fränkl during lauge koch’s 1950–51 expeditions to the present nordsylen, a mountain about 1500 m high in the northern stauning alper. the name appears on the profiles in fränkl (1953). girl’s name. birkedal 75ø-80 (75°00.1´n 22°10.3´w; map 4). valley on the west side of inner grandjean fjord. the name originated from the wintering party at kulhus in 1935, and was given for the occurrence of the dwarf birch. see also betula havn. birkedalshytten 75ø (75°01.1´n 22°03.5´w). alternative name for betulahavnhytten, a danish hunting hut built in 1951 at betula havn, inner grandjean fjord. it is sited at the mouth of birkedal. bischofsmütze 72ø (72°07.5´n 25°33.7´w). mountain 1360 m high in the trekantgletscher area, west of alpefjord. climbed by wolf gang weinzierl’s 1970 expedition, and named after the austrian peak of the same name, the highest mountain in the dachstein group. (bishop’s mitre.) fig. 32. looking south at part of berzelius bjerg, the spectacular mountain formed in multicoloured late precambrian rocks of the eleonore bay supergroup; north side of segelselsällskapet fjord, lyell land. 133 bishops glacier 72ø (72°22.6´n 25°23.7´w; map 5). name used by bennet (1972) for a glacier in ne nathorst land draining east to alpefjord. biskop alfs gletscher 71ø (71°51.1´n 24°02.7´w). glacier draining the south flank of the werner bjerge, the present breithorn glet scher. the name was one of a group of names for glaciers given by the place name committee in 1939, which replaced proposals by hans stauber. alf (álfr) was bishop to the norse settlers of green land from 1365 to 1378. the name was officially approved from 1939 to 1956, and appears on some later published map sheets, but has rarely been used in scientific reports. in 1956 the name was formally abandoned in favour of the more commonly used name breit horn gletscher. biskop joseph fjeld 71ø (71°07.0´n 21°53.6´w). the name is used in den grønlandske lods (1968) for a mountain in liverpool land, the present kirken. it was one of the names introduced by henning bistrup on his coast profiles drawn in 1923 and 1930. bison lake 77ø (77°35.3´n 20°48.8´w). lake south of klægbugt, nordmarken. named by the 1987 irish expedition to ne green land. bispehuen 71ø-389 (71°34.4´n 23°35.6´w). mountain 1261 m high east of pothorst bjerge, northern jameson land, with a shape said to resemble a bishop’s mitre. the name was suggested by russel marris following his explorations in 1968. bivuakkammen 71ø-245 (72°00.2´n 23°54.9´w). ridge in the werner bjerge. named by peter bearth and eduard wenk during lauge koch’s 1953 expedition for the locality of one of their satellite camps used during geological exploration (kam = ridge). bjarne larsenfjellet 74ø (74°18.1´n 20°48.3´w). mountain 1293 m high on central clavering ø. used only on nsiu maps (lacmann 1937), the name was given for bjarne larsen [b.1904], a norwegian aeroplane mechanic who took part in the 1932 nsiu expedition. bjergandesø 76ø-239 (76°49.1´n 19°08.4´w). lake on winge kyst in southern germania land. so named by the 1906–08 danmarkekspeditionen because a pair of scaups (= bjergænder) were ob served here in june 1907, an uncommon bird in this part of east greenland. bjergkronerne 71ø-163 (71°51.5´n 23°31.1´w; map 4). mountain range in scoresby land north of ørsted dal, with summits formed by basalt sills giving a crown-like appearance (krone = crown). named during lauge koch’s 1936–38 expeditions by hans stauber. bjerring pedersen fjæld 70ø (70°30.6´n 22°08.8´w). name used by rosenkrantz (1934, 1942) after one of the summits of gulfjelde in southern liverpool land, an area first investigated by bjerring pedersen in 1924. see also bjerring pedersen gletscher. (mt. bjerring pedersen.) bjerring pedersen gletscher 70ø-231 (70°43.9´n 21°53.3´w). glacier in southern liverpool land draining east to the head of vejle fjord. named by laurits bruhn during the 1931–34 treårs ekspeditionen after bjerring pedersen [1898–1925], a danish geologist who led the party of six scientists and carpenters during the 1924–25 colonisation expedition to scoresbysund, until his death on 2 july 1925. (bjerring pedersens gletscher.) bjørktun 72ø (72°24.4´n 26°02.7´w). norwegian hunting hut built in september 1931 by the møre expedition, about 10 km from the inner end of forsblad fjord. it was destroyed by an avalanche in the spring of 1976. the hut was named after the small birch trees surrounding the hut (rogne 1981), and has also been known as inderhytten. bjørn jorsalfarers gletscher 71ø (71°57.0´n 24°43.0´w). glacier in the stauning alper draining east to schuchert dal, the present storgletscher. the name was one of a group of names for glaciers given by the place name committee in 1939, which replaced proposals by hans stauber. the name was officially approved from 1939 to 1971, although rarely used on maps. it commemorated the widely travelled icelander björn einarsön of vatnafjord, also called björn jorsalsfarer (or jorsalfarer), who in 1385 made a voyage to greenland. due to confusion arising from inaccurate topographical maps the names langgletscher and storgletscher were also applied to the same glacier, until in 1971 storgletscher became the only approved name. bjørn pynt 76ø-81 (76°37.6´n 18°35.9´w). cape on eastern lille koldewey, so named by the 1906–08 danmark-ekspeditionen, possibly after poul harald bjørn who at one time worked with bendix thostrup at the danish nautical charts archive (j. løve, personal communication 2009). (björns pynt, bjørn odde.) bjørnbo gletscher 71ø-153 (71°40.1´n 24°53.8´w; maps 4, 5). large glacier in the southern stauning alper draining se into schuchert dal. the name was one of a group of names for glaciers given by the place name committee in 1939. it commemorates axel anthon bjørnbo [1874–1911], a noted authority on green land and author of ‘cartographia groenlandica’. (bjørnbos glet scher.) bjørnbos corner 71ø (71°41´n 24°30´w). name used by kempter (1961) for the area ne of the terminus of bjørnbo gletscher where it meets schuchert dal; this has become the type locality of his ‘bjørnbos corner formation’. the name is not approved. bjørnbos elv 71ø (71°37.5´n 24°34.2´w). name occasionally used in reports of the 1962 oxford university expedition for the river draining bjørnbo gletscher. bjørnebu 72ø (72°07.5´n 23°28.6´w). norwegian hunting hut built in august 1930 for the møre expedition east of the mouth of mesters vig. it was originally called jostein, and has also been known as pictetbjergshytten and segldalen. bjørnebu 74ø (74°24.2´n 19°09.5´w). norwegian hunting hut on the coast of wollaston forland, on the north side of dronning augustadalen. built by the hird expedition in july 1928, it has also been known as stordalen and augustadalshytten. hunters considered it to be one of the best huts on the coast (p.s. mikkelsen 1994). bjørnedal 72ø-102 (72°20.9´n 22°47.3´w). valley on se traill ø between mountnorris fjord and kong oscar fjord. so named during the 1931–34 treårsekspeditionen by ove simonsen be cause while surveying here two bears were shot. the valley appears to be a bear migration route between the two fjords. the names barrikadental and volldal have also been used for this valley. bjørnedal 73ø (73°20.3´n 22°14.5´w). valley in the southern giesecke bjerge, corresponding to the present randbøldalen. so named on the nsiu (1932a) map, and given for the polar bear. franklindalen has been used for the same feature (bang 1944). bjørnegletscher 80ø-58 (80°36.3´n 17°56.5´w; map 4). glacier on the west side of outer ingolf fjord, draining the prinsesse elisabeth alper. named by the 1938–39 drastrup/kristoffersen expedition, possibly for the incident recorded by kristoffersen (1969) of a bear that had followed their ski tracks up to the sledge and ‘insisted’ on being shot (fig. 33). bjørn-heimen – see borganes. bjørneheimen 73ø (73°07.6´n 25°44.4´w). norwegian hunting hut on the west side of antarctic sund at the mouth of nanortalikdal, andrée land, built in september 1934 by arktisk næringsdrift. this hut is often referred to as nanortalik or nanortalikhytten. bjørneheimen (= home of the bear) has almost the same meaning as nanortalik (= the place where there are many bears). (bjørne heimhytta.) bjørnehytten 73ø (73°26.9´n 20°38.0´w). norwegian hunting hut built for arktisk næringsdrift in september 1950 at kap broer ruys, the se point of hold with hope. so named because a bear had disturbed the depot of materials intended for the hut, and dragged a complete wall out to sea. the wall was replaced by planks intended for the floor (p.s. mikkelsen 1994). the hut is also known as gnisten. bjørnepas 72ø-165 (72°28.8´n 22°15.2´w). pass between æbeltoft 134 vig and begtrup vig, eastern traill ø, a migration route for bears. named during lauge koch’s 1936–38 expeditions by hans peter schaub. bjørneskærene 77ø-41 (77°50.8´n 19°14.7´w). line of skerries ne of gamma ø. so named by the 1906–08 danmark-ekspeditionen because they encountered many bears here, of which two were shot. (björneskærene, bear skerries, bjørne skærene, bjørneskær). bjørnesø 72ø-228 (72°08.0´n 23°43.4´w). small lake on the east side of mesters vig. named by prospecting teams associated with lauge koch’s 1948–49 expeditions, for an incident involving a bear. bjørnesø 72ø-240 (72°51.8´n 25°06.2´w). lake se of lauge koch’s scientific station on northern ella ø. the name is said to have originated in the 1930s with aage de lemos, lauge koch’s telegraphist at the station, and was given for an incident involving a bear. (bear lake). bjørnesø 76ø-165 (76°08.5´n 20°30.1´w). lake in påskedalen, ad. s. jensen land. the name first appeared on the map of the 1932 gefion expedition (jennov 1935), and was so named because on his first visit to the lake j.g. jennov had seen bear tracks. bjørnesø 72ø (c. 72°13´n 23°54´w). lake near mesters vig where samples were taken for radiocarbon age determinations (cremer et al. 2008). bjørnetoppen 74ø (c. 74°17´n 19°25´w). name used by the 1908–09 floren expedition for a hill near their house at kap borlase warren, se wollaston forland, where they shot bears in october and december 1908. the name is used in the diary of the expedition published by brandal (1930), although the exact position is uncertain. bjørneø 73ø-516 (73°33.4´n 24°44.1´w). small island near the mouth of geologfjord, so named by a.g. nathorst in 1899 as björnön because a polar was shot here by ivar arwidsson, the expedition zoologist. (bear island, bjørnøn.) bjørneøer [nannut qeqertaat] 71ø-42 (71°07.0´n 25°25.0´w; maps 3, 4; fig. 34). group of islands off ne milne land. so named by carl ryder’s 1891–92 expedition as the bjørne øer, because a bear was shot during exploration of the islands on 4 september 1891. the main islands are sometimes numbered i to xi (see e.g. kalsbeek 1969), the notation deriving from the first survey carried out by eduard wenk and helge backlund in 1934. the 1934 party climbed parts of the spectacular ridges of islands vi and ix (första nålbrevet, sista nålbrevet), while further climbs were made in 1978 by a british army expedition. (björneöer, bjørne islands, bear islands, bären inseln.) bjørnnesstua 74ø (74°27.1´n 21°41.9´w). norwegian hut se of the giesecke bjerge built in july 1932 by the w. holmboe salmon fishing expedition. it has also been known as holmboehytten and giskehuset. bjørnselv 74ø-201 (74°16.2´n 20°26.5´w). minor stream on eastern clavering ø, between storstrømmen and grønnedal, so called by danish hunters. the name first appeared on a sketch map in gustav thostrup’s 1921 logbook (møller 1939). the name has apparently also been used for the present henningselv, and on some ams maps has been applied to the river in grønnedal. black hills 73ø (73°18.7´n 25°03.7´w). area of low hills between noa sø and innermost dusén fjord, ymer ø. the name was given for the colour of the rocks by cleaves & fox (1935) during geological work on the 1933 john k. howard expedition. black mountain 70ø (70°47.0´n 25°58.6´w). mountain 1635 m high south of korridoren, milne land, carved into black basaltic lava flows. climbed by the 2004 west lancashire scouts expedition, after an approach by ski. black twin – see schwarzer zwilling. blackwall 72ø (72°10.0´n 24°39.5´w; map 5). mountain 1850 m high between dunottar gletscher and harlech gletscher, northern stauning alper. first climbed by the 1963 imperial college expedition and named, like many of their other ascents, after a london locality. blackwall is a district of poplar on the north bank of the river thames. blair 71ø (71°42.8´n 25°20.0´w; map 5). mountain about 2200 m high on the sw side of orion gletscher, southern stauning alper. first climbed by james clarkson’s 1961 expedition, and named after blair castle, a scottish mansion built by the duke of atholl in 1269, and rebuilt in 1869. blanke bugt 70ø-410 (70°32.5´n 25°56.5´w). small bay on sw milne land, sw of mudderbugt. named during the 1967–72 ggu scoresby sund expeditions by svend funder for the calm water. blaskbjerg 73ø-596 (73°18.6´n 24°01.7´w). mountain in gunnar andersson land, northern ymer ø. the name was used by gunnar säve-söderbergh during the 1931–34 treårsekspeditionen in the form mt. blask, because he had expected to find fossils here, and when he didn’t it was a ‘blask’ (blask = splash, which can best be fig. 33. bjørn(e) or strictly isbjørn (polar bear) is a common visitor in east greenland fjords. this mother with two large cubs was photographed from a cruise ship in the pack ice off the coast. photo: adam a. garde. 135 translated here as ‘a flop’). the name was said to have been introduced by aage de lemos, one of the wintering party in 1931–32 on ella ø. blastfjord – see føhnfjord. blattspitze 72ø (72°08.2´n 25°42.1´w). mountain 2000 m high in the trekantgletscher area, west of alpefjord. climbed and so named by wolfgang weinzierl’s 1970 expedition. exact location a little uncertain. (leaf peak.) blika 73ø (73°37.4´n 21°52.1´w). river flowing into the south end of loch fyne. so named on the nsiu (1932a) map, and apparently derived from a norwegian dialect word for a white stripe on a hillside. blindeskær 71ø-47 (71°47.2´n 22°13.6´w). submerged rock 1.5 km off kap tyrrell, the nw point of canning land, which the ant arctic sailed over on 24 august 1900. named by g.c. amdrup’s 1898–1900 expedition. blindtarmen 71ø-109 (71°08.9´n 21°50.8´w). short fjord in east liverpool land nw of kap jones. so named by laurits bruhn during the 1931–34 treårsekspeditionen (blindtarm = appendix). blindtarmen 74ø (74°34.3´n 22°00.0´w). name sometimes used by norwegian hunters for the narrow nw–se-trending inner part of tyrolerfjord, between payer land and a.p. olsen land. blindtarmen 79ø-22 (79°47.7´n 19°45.2´w). deep bay on the east side of inner dijmphna sund. the name is attributed to david malmquist, and arose during lauge koch’s geological expeditions. it was approved in 1958. bloch nunatakker 79ø-12 (79°37.1´n 20°29.6´w; maps 1, 4). nunatak group on the north side of lambert land, named by the 1909–12 alabama expedition after commander bloch of the hekla, who had assisted the expedition. the place name com mittee position for this group of nunataks, adjacent to the north point of lambert land, is probably incorrect. from his position high on the inland ice it is more likely that ejnar mikkelsen observed the small group of islands that split the ice front filling nioghalvfjerdsfjorden at about 79°37´n 20°29´w. jacob christian demant bloch [1859–1944] was commander of the cruiser hekla and had helped ejnar mikkelsen with his sick dogs in the faeroe islands (j. løve, personal communication 2009). (bloch’s nuna takker.) blockfjellet 74ø (74°19.8´n 21°17.1´w). mountain on central clav ering ø, named on nsiu maps (lacmann 1937) after walter block [b. 1902], who assisted in the photogrammetric construction of the fig. 34. original survey of bjørneøer, the island group south of the mouth of nordvestfjord, drawn by helge backlund and eduard wenk in 1934. the sketch map shows sight lines from trigonometric points (tp) and the numbering system in roman numerals. geus archive. 136 nsiu maps of east greenland. blokadedal 73ø-61 (73°43.7´n 22°35.3´w). valley in hudson land west of stordal. so named by lauge koch’s 1929–30 expeditions in the form blocade valley, because the mouth of the ice-filled valley is partially blocked by moraine. (blockade tal). blokdal 73ø-50e (73°58.3´n 21°24.8´w). minor valley in nw hold with hope on the north slope of stensiö plateau, draining into blåelv. so named during the 1931–34 treårsekspeditionen by eigil nielsen, probably because of the numerous fossiliferous boulders. blokelv 70ø-299 (70°29.5´n 22°07.9´n; map 4). river in south liverpool land west of scoresbysund [illoqqortoormiut]. so named by alfred rosenkrantz during lauge koch’s 1926–27 expeditions as blok elv or block river, because in 1926 he found here two large boulders containing a particular suite of fossils. blokelv 70ø-97 (70°40.5´n 23°51.7´w). river in jameson land flowing sw to enter the sea close to vandreblokken. so named by laurits bruhn during the 1931–34 treårsekspeditionen, after vandreblokken. blokelv 73ø-50d (73°58´n 21°25´w). minor river in blokdal, on the north slope of stensiö plateau, draining into blåelv, nw hold with hope. named during the 1931–34 treårsekspeditionen by eigil nielsen. blokken 73ø-597 (73°48.0´n 24°32.3´w; map 4). mountain in eastern strindberg land. the name was first used by teichert (1933) during the 1931–34 treårsekspeditionen (blokken = the block). blokken 74ø (74°01.7´n 21°37.0´w). name occasionally used by eigil nielsen for knolden, a minor feature north of frebold bjerg, hold with hope. blomsterbjerg 73ø-401 (73°27.7´n 25°57.0´w). mountain in southern andrée land between luciagletscher and blåbær glet scher, named by erdhart fränkl during lauge koch’s 1948–50 expeditions for the abundant flowering plants (= blomster). blomsterbugten 73ø-562 (73°19.7´n 25°17.5´w; map 4; fig. 35). bay on the west coast of ymer ø. so named by gunnar seidenfaden and arne noe-nygaard during lauge koch’s 1929 expedition be cause abundant botanical collections were made here. vargbukta has also been used. (blomster bay, blomster bukta, bay of flowers.) blomsterbugthytten 73ø (73°19.9´n 25°16.9´w). norwegian hunting hut built for arktisk næringsdrift in march 1930 at blomster bugten, west ymer ø. it was originally known as varghytta. (blom ster bugt.) blomsterdal 71ø-254 (71°57.8´n 23°36.9´w). valley in northern scoresby land draining via kolledalen into antarctic havn. the name was given by the place name committee about 1956 as a replacement for a suggestion by peter bearth. (blomsterdalen.) blomsternunatak 72ø-289 (72°44.2´n 28°08.4´w). large nunatak on the north side of hisinger gletscher, west of gletscherland. named during lauge koch’s 1953 expedition by john haller, for the abundant flowers. blomstersø 72ø-481 (72°20.6´n 26°16.5´w). lake at the nw end of snedrivegletscher, sw of tærskeldal, nathorst land. named during lauge koch’s 1954–55 expeditions by hans zweifel, for the many flowers along its shores. blosseville bjerg 74ø-79 (74°15.7´n 22°11.1´w; map 4). mountain 1283 m high west of clavering ø. karl koldewey’s 1869–70 expedition had given the name cap blosseville in commemoration of jules baron de blosseville [1802–33], a lieutenant in the french navy who disappeared without trace off the east coast of greenland in the la lilloise in 1833. the name appeared in the form cape blosseville on the maps of j.m. wordie’s 1926 expedition and lauge koch’s 1929–30 expeditions, and also on nsiu maps (lacmann 1937). it was transferred to the mountain by the place name com mittee about 1934 because of discrepancies between koldewey’s description and map and modern maps. koldewey’s map does not show granta fjord, and he appears to have mistaken the present blosseville bjerg for jordanhill. (cape blosseville, kapp blosseville.) blosseville kyst 68ø-17 (69°00´n 26°00´w; map 3). name in general use for the inhospitable coastal stretch of basalt cliffs extending from about 68°n to 70°10´n. officially the name applies only to that part of the coast from 68°–69°n surveyed by jules baron de blosseville [1802–33] in 1833, but it is nearly always used in a wider sense (e.g. in den grønlandske lods, 1968). on early maps the coast was marked land opdaget af j. de blosseville or simply blosse ville 1833, and appears first as blossevilles kyst on the maps of the 1879 ingolf expedition. blosseville was a french marine officer who had made several voyages to the west indies, south america, india and burma, and was lost with his ship the la lilloise and his entire crew on this stretch of coast in 1833 (j. løve, personal com munication 2009). see also blosseville bjerg. (blossevilles-kyst, de blosseville coast, côte de blosseville.) bluie east 3 71ø (71°14.7´n 24°35.0´w). code name used by the us coast guard during the second world war for gurreholm, on the west coast of jameson land. lt. arnold peterson of the us coast guard wintered at gurreholm in 1941–42. the code names beginning bluie west were in west greenland. bluie east 4 71ø (72°52.6´n 25°06.7´w). code name used by the us fig. 35. view over noa sø on ymer ø across blomsterbugten and central kejser franz joseph fjord to hvidevæggen in andrée land, with payer tinde in suess land in the background. the john haller photograph collection, geus archive. payer tinde suess land kejser franz joseph fjord blomsterbugten noa sø hvidevæggen ymer ø 137 coast guard during the second world war for ella ø station, ella ø, built by lauge koch in 1931. bluie east 5 74ø (74°05.7´n 21°16.8´w). code name used by the us coast guard during the second world war for eskimonæs, lauge koch’s scientific station on south clavering ø. the same code name was used for myggbukta after eskimonæs was destroyed in 1943. blyklippen 72ø-188 (72°11.2´n 24°07.2´w; map 4; fig. 36). hill side to the west of store blydal, north scoresby land. named by prospecting teams associated with lauge koch’s 1948–49 expeditions for the presence of lead ore (bly = lead). a lead mine, sometimes referred to as blyklippen mine, was excavated beneath bly klippen in a major quartz vein containing a sulphide lens, and between 1956 and 1962 yielded 545 000 tons of lead-zinc concentrate (a. mikkelsen 1992; thomassen 2005a). blyryggen 72ø-213 (72°08.9´n 23°56.6´w; map 5). ridge west of the bay mesters vig, north scoresby land, rising to 1051 m. named by prospecting teams associated with lauge koch’s 1948–49 expeditions for finds of lead ore. blæsebræ 78ø-35 (78°12.9´n 21°23.4´w; maps 1, 4). glacier between søndre mellemland and nørre mellemland, hertugen af orléans land. named by the 1938–39 mørkefjord expedition, for the strong katabatic winds. blæsebælgen 74ø (74°32.2´n 18°48.3´w). name used by hvidberg (1932) for the hut in germaniahavn on sabine ø where he experienced a violent two-day storm in august 1929 (blæsebælgen = the bellows). the hut has usually been known as germaniahavn or villaen. blæsedal retrætdepot 74ø (74°16.9´n 19°51.7´w). hut built by nordøstgrønlands slædepatrulje in december 1944 about 3 km into blæsedalen, wollaston forland. it was intended to support a retreat position should the main station at daneborg be attacked by german forces. blæsedalen 74ø-244 (74°18.4´n 19°49.0´w). valley west of her schell bjerg, wollaston forland. it records the strong katabatic winds. glesdalen has also been used. blæsedalhytten 74ø (74°21.8´n 19°47.6´w). danish hunting hut built by nanok in may 1947 about 11 km up in blæsedalen, wollaston forland. blæsedalhytten 74ø (74°23.6´n 19°46.7´w). norwegian hunting hut built in may 1947 by hermann andresen’s expedition about 15 km up in blæsedalen, wollaston forland. this hut was extended by sirius in 1961 (p.s. mikkelsen 1994). blæsenborghytten 74ø (74°30.2´n 20°37.9´w). danish hunting hut at the east end of store sødal, ne of zackenberg. built by nanok in august 1938, it was named by one of the hunters for very strong winds experienced in november 1938. it has also been known as dalhytten. blödelbreen 74ø (74°20.3´n 21°18.1´w). glacier on central clav ering ø. used on the nsiu maps of lacmann (1937), the name is derived from the german epic poem from c. 1200, the nibelungen lied. blå sø 70ø (70°15.4´n 28°58.1´w). name reported used by catalina pilots in 1958 for the present kaskadesø in western gåseland. eduard wenk noted that it was an inappropriate name as the lake is grey rather than blue in colour. blaabær-cliff 74ø (74°50.2´n 20°44.7´w). name used as a geological reference locality by wolf maync (1947) for the cliff section south of the danish hunting hut, blåbærhytten, on the west side of fligely fjord. the danish ‘blåbær’ (= blueberry) are equivalent to the edible bilberry. (blaabærcliff, blaabærhytten-cliff.) blaabærdal 75ø (75°15.4´n 21°03.7´w). name used by danish hunters for the present kildedal, a valley on the south side of ardencaple fjord (hvidberg 1932). it was named for its berries. blåbærdal 73ø-400 (73°23.8´n 25°40.5´w). valley in southern andrée land. named during lauge koch’s 1948–50 expeditions by erdhart fränkl for its berries. (blaabærdal, blaabær dal.) blåbærdalen 74ø-297a (74°52.0´n 21°06.4´w; map 4). valley in east th. thomsen land, draining into fligely fjord. named probably by nanok, which built a hut at the mouth of the valley in 1931. berries are plentiful here in the summer. the original form was blaabærdal. (blaabær dal.) blåbærfjeldet 76ø-150a (76°40.2´n 21°40.1´w). mountain in southern daniel bruun land, named by j.p. koch during his 1912–13 expedition as blaabærfjældet. he found large quantities of bilberries here towards the end of a strenuous 35-hour march in 1912. blåbærgletscher 73ø-399 (73°28.8´n 25°47.5´w; map 4; fig. 37). fig. 36. the entrance to the closed lead mine at blyklippen, near mestersvig. the mine was worked from august 1952 until may 1963. 138 glacier in southern andrée land draining via blåbærdal to eleo nore bugt. named during lauge koch’s 1948–50 expeditions by erdhart fränkl. (blaabær gletscher.) blåbærhus – see blåbærhytten. blåbærhytten 74ø-297 (74°50.3´n 20°44.3´w). approved name for the danish hunting hut on the west side of fligely fjord at the mouth of blåbærdalen built by nanok in august 1931. nyholmpoulsen (1985) described it in 1932 as a simple hut, two by two metres, with a roof of musk-ox skin. it was extended in august 1951, and unofficially promoted to blåbærhus. (blaabærhytten.) blåbærhøj 70ø-64 (70°27.4´n 26°14.6´w). small hill about 80 m high north of hekla havn on danmark ø. so named by carl ryder’s 1891–91 expedition as blaabærhøjen, because bilberries were common here. blåelv 73ø-47 (73°58.7´n 21°21.9´w). river in home forland draining north into godthåb golf. named by lauge koch’s 1929–30 expeditions in the form blue river, for the occurrence of bluish grey sandy shales of carboniferous age. the name is found in koch (1931), and corresponds to his river 16. it may be the same as that originally named wordie creek by lauge koch, which has been identified with either river 15 or river 16; see also wordie kløft. (blåelva, blaaelv.) blåhorn 72ø-244 (72°20.5´n 24°43.2´w; map 5). mountain 1589 m high in the northern stauning alper. named by erdhart fränkl during lauge koch’s 1950–51 expeditions, for the colour of the rocks. (blaahorn.) blåhö ( full name = blåhögda) 73ø (73°35.3´n 21°17.5´w). mountain 1067 m high, part of the present ravnebjerg, hold with hope. so named on an nsiu map (nsiu 1932a; fig. 13). blårævekløft 73ø-667 (73°18.3´n 26°03.6´w). minor ravine draining into junctiondal, south andrée land. named during lauge koch’s 1948–50 expeditions by erdhart fränkl, after the blue fox (= blåræve), one of which stole some of fränkl’s underwear. (blaarævekløft.) blåræven 74ø (c. 74°11´n 22°13´w). norwegian hunting hut built in 1935 for arktisk næringsdrift about 3 km ne of hansen havn, at the front of wordie gletscher; now disappeared. it was also known as hansen havnhytten. blåsærk 69ø (69°03.0´n 26°49.3´w). mountain behind the blosse ville kyst, equivalent to the present rigny bjerg according to tornøe (1935, 1944). the name features in several of the icelandic sagas (landnámabók, eirik raudes saga, torfinn karlsevnes saga), and is usually given in the form bláserkr (blåsærk = blue shirt). other authorities have located this feature farther south in se greenland. blåserkjøkulen 69ø (69°00.0´n 26°34.0´w). in the icelandic sagas blåserk is used both for the mountain and the glacier from which the mountain rises. tornøe (1935) has argued convincingly that blåserk is identical with rigny bjerg, as seen from the sea; rigny bjerg rises from behind a marked glacier that he terms blå ser k jøkulen. blåsø 79ø-9 (79°35.0´n 22°30.0´w; maps 1, 4). tidal lake in southernmost kronprins christian land, dammed by the floating glacier filling nioghalvfjerdsfjorden. it was mapped from the air by lauge koch during the 1931–34 treårsekspeditionen and named for its blue colour. (blaasö lake.) bocksrietdalen 72ø-417 (72°53.8´n 27°33.4´w). broad valley at the head of kjerulf fjord extending southwards to hisinger glet scher. it was named during the 1931–34 treårsekspeditionen by eugène wegmann as bocksriet valley, after a locality near schaff hausen in switzerland. in swiss german dialect the name means a place ‘where the devils are dancing’ (fritz schwarzenbach, personal communication 1996). new valley and kjerulfsdalen were used by louise boyd, who explored the valley in 1931 and 1933. bodal 70ø-153 (70°51.8´n 22°23.1´w). valley in liverpool land on the east side of hurry inlet. named during the 1931–34 treårs ekspeditionen by laurits bruhn. bodger 71ø (71°09.3´n 26°36.9´w). summit 1954 m high on the ice cap between catalinadal and edward bailey gletscher, renland. climbed and named by the 2007 west lancashire mountaineering group expedition. bohr bjerg 77ø-129 (77°04.7´n 24°47.7´w; map 4). prominent mountain in nw dronning louise land. one of the names given by the 1952–54 british north greenland expedition for notable blåbærgletscher andrée land fig. 37. looking north-west across the glaciers and alpine mountains of andrée land. the folded rocks on the north side of blåbærgletscher were deformed during the caledonian orogeny. the john haller photograph collection, geus archive. 139 scientists, it was named after the danish physicist and nobel laureate niels bohr [1885–1962]. he made major contributions to the development of quantum physics, and was responsible for the bohr theory of the atom. bolettestua 74ø (74°57.3´n 20°02.4´w). norwegian hunting hut built in august 1932 for sigurd tolløfsen’s expedition about 5 km south of kap bremen, kuhn ø. it was named after tolløfsen’s wife bolette. boksehandsken 71ø (70°44.1´n 24°03.1´w). lake in western jameson land about 54 m above sea level. this informal name is used in descriptions of the quaternary geology of the area (ingólfs son et al. 1991), and reflects the shape of the lake that resembles a boxing glove (= boksehandske). (lake boksehandsken.) bolværket 72ø-322 (72°04.8´n 24°59.7´w; map 5). mountain 2571 m high on the south side of gullygletscher, north stauning alper. named during lauge koch’s 1954 expedition by john haller, for its appearance (bolværk = bulwark). it was first climbed by the 1963 cambridge university expedition, that considered it their most difficult peak of the summer. bonar bjerg 71ø (71°51.7´n 24°52.1´w; map 5). mountain 2241 m high between gannochy gletscher and roslin gletscher, south stauning alper. first climbed by the 1968 university of dundee expedition, and possibly named after bonar bridge, a village on the kyle of sutherland, scotland. bonney plateau 73ø-338 (73°26.3´n 22°04.8´w). flat-topped mountain in the central giesecke bjerge. it commemorates tho mas george bonney [1833–1923], a british theologian and naturalist noted for his popular texts on geology. the mountain was climbed by noel odell and walter wood in 1933. breidhausen has also been used. (bonneys plateau.) bonsachs ø 76ø (76°45.0´n 20°41.8´w). island east of daniel bruun land, the present ringøen. so named by the 1932 gefion expedition. bontekoe ø 73ø-4 (73°07.0´n 21°22.5´w; maps 3, 4). large island in foster bugt. the name occurs on charts published in paris by f.e. foster in 1783 and 1788, and that published in 1818 to accompany hidde dirks kats ‘dagboek eener reize i de jaren 1777 en 1778’. the name was adopted by william scoresby jr. in 1822 as bontekoe island, but he incorrectly identified it with kap broer ruys, and his own cape humboldt was probably the real bontekoe ø. the island was correctly placed se of kap franklin on clav ering’s (1830) maps. the name is probably that of the dutch whaler who first sighted the island. a.g. nathorst built a cairn on the summit in 1899, his message being retrieved by adolf hoel in 1931. (bontekoe ö, bonte koe ön, bonte-koe-insel, bontekoes ø, bonteko ø, bonteco.) bopladsdalen 70ø-354 (70°07.3´n 22°14.0´w). valley behind the settlement at kap brewster. the name was used by hassan (1953) in his report on material collected during lauge koch’s 1951 expeditions (boplads = settlement). bordbjerget 72ø-230 (72°44.9´n 23°19.0´w). mountain on ne traill ø, a flat-topped summit in the rold bjerge. so named by desmond donovan during lauge koch’s 1949–50 expeditions (bord = table). bordet 74ø-385 (74°06.5´n 29°03.5´w). ice-covered area at the margin of the inland ice nw of hobbs land. this name was given by the place name committee replacing a suggestion by hans r. katz. the area is flat and notable for a lack of crevasses (bordet = the table). boreal zenith 73ø (73°32.5´n 26°10.1´w). impressive summit with fine views on the south side of grejsdalen, andrée land. climbed by the 2007 army boreal zenith expedition. borestok 72ø-399 (72°03.5´n 23°30.9´w). ridge in northern scores by land between jægerdal and segldal. so named by hans kapp during lauge koch’s 1957–58 expeditions because of the three stake-like humps on the ridge eroded in basalt (bore = drill, stok = stick, stake). boresø 74ø (74°30.4´n 20°37.3´w). lake in the zackenberg area where samples were taken for radiocarbon age determinations (cremer et al. 2008). borg 76ø (c. 76°42´n 22°24´w). wintering station of j.p. koch’s 1912–13 expedition in the centre of bredebræ, to which it had been transported by pony and boat from danmarkshavn. the name had been given by koch’s wife after eigil skallegrimsson’s farm, borg, in iceland (p.s. mikkelsen 1994). the station was only manned in the winter of 1908–09, and subsequently disappeared, carried out into dove bugt by the calving of the glacier. borganes 74ø (74°15.9´n 19°22.9´w). primitive norwegian hunting hut at kap borlase warren, wollaston forland, originally built in 1908 by severin liavaag’s floren expedition on the ruins of an inuit house. the old hut and the cape have been known by a variety of names: bjørn-heimen, gammen, sverdrupsnes (p.s. mikkelsen 1994); see also grønlænderhuset. østgrønlandske fangstkompagni built a house at the same site in 1922 known as valdermarshaab, which was taken down in 1923 following a poor trapping season. the hird expedition repaired the old hut in 1927. in his diary of the 1908–09 expedition brandal (1930) employs this name for the cape itself, which as giæver (1958) notes was entirely appropriate as the cape resembles a stone castle (= borg). (borgarnes.) borgbjerg gletscher 71ø-61 (71°40.0´n 25°50.0´w). major glacier on the north side of central nordvestfjord. one of the names used on the 1932 edition of the geodetic institute 1:1 million scale map, it derives from lauge koch’s aerial observations during the 1931–34 treårsekspeditionen. frederik borgbjerg [1866–1935] was a member of the danish parliament from 1892, and minister of education from 1929. he was present at the departure of treårs ekspeditionen from copenhagen in 1931. borgbjergkamm 71ø (71°48.1´n 25°44.6´w). ridge up to 2400 m high leading northwards to borgbjergtinde, in the ne part of the borgbjerg gletscher region, southern stauning alper. probably named by the 1977 schwäbische stauning alper expedition. borgbjergtinde 71ø (71°49.6´n 25°43.5´w; map 5). peak 2546 m high in the ne part of the borgbjerg gletscher region, southern stau ning alper. probably first climbed and named by the 1977 schwäbische stauning alper expedition. borgen 70ø-264 (70°06.0´n 23°42.4´w; map 4). mountain on volquaart boon kyst flanked by østre borggletscher and vestre borggletscher. named during the 1931–34 treårsekspeditionen by laurits bruhn for its castle-like appearance. borgen 74ø-228 (74°01.3´n 21°34.2´w). feature in nw hold with hope, named by eigil nielsen during the 1931–34 treårs ekspe ditionen. the position is said to be uncertain because it is not found on his map, but from the description in the text (nielsen 1935 p. 49), it is most probably the ridge between pyramiden and fiskeplateau. borgfjorden 76ø-116 (76°40.0´n 22°00.0´w; maps 2, 4). fjord between daniel bruun land and lindhard ø, so named by j.p. koch during his 1912–13 expedition because it lay east of the wint ering station borg. (borg fjorden, borg-fjord, castle fjord, borgar fjorður.) borggletscher 73ø-439 (73°02.9´n 26°40.1´w). glacier in suess land between østre vikingeborg and vestre vikingeborg. named by john haller during lauge koch’s 1952–53 expeditions. borgjøkelen 76ø-126 (76°38.5´n 23°48.0´w; map 2). glacier in central dronning louise land, so named by j.p. koch during his 1912–13 expedition as borgjøklen or borgjöklen, because it was situa ted west of the wintering station borg. the definitive ‘en’ end ing was part of the approved name for many years, but has been omitted on the most recent lists of authorised names. (borgjøkelen, borgarjökull.) borgvig 70ø-261 (70°08.5´n 23°51.9´w). bay at the front of vestre borggletscher, west of borgen, volquaart boon kyst. named during the 1931–34 treårsekspeditionen by laurits bruhn. 140 borgøen 73ø-247 (73°05.9´n 22°34.3´w; map 4). largest island in the broch øer group, east of ymer ø. the name seems to appear first on an nsiu map (nsiu 1932a) in the form borgöya, and was presumably named for a castle-like appearance. bórje elv – see börje elv. borlase warren hytten 74ø (74°15.9´n 19°23.0´w). danish hut built by sirius in the summer of 1956 at kap borlase warren, wollaston forland (p.s. mikkelsen 1994). bosigran 72ø (72°08.1´n 24°54.9´w; map 5). pinnacle about 2700 m high on the ne ridge of hjørnespids, north stauning alper. climbed by the 1968 queen mary college expedition on 13 august, and named after a climbing locality in cornwall. botanikerbugt 73ø-595 (73°02.3´n 24°39.2´w; map 4). bay on the south coast of ymer ø. the name was used as a botanical reference locality in reports of the 1931–34 treårsekspeditionen. thorvald sørensen carried out detailed botanical studies here. (botaniker bugt). bothriolepis cleft 73ø (73°35.3´n 23°52.2´w). ravine on the south side of gauss halvø, west of paralleldal. the name was used by gun nar säve-söderbergh during the 1931–34 treårsekspedi tion en, because of finds of fossil bothriolepis (säve-söderbergh 1934). bothriolepis mtn 73ø (73°22.0´n 24°11.0´w). name used by stensiö (1936) for a mountain on the north side of ymer ø where de vonian fossils (bothriolepis) were collected in 1934. location uncertain, but it may be the 826 m high mountain east of the mouth of zoologdal. botnhuset 73ø (73°40.6´n 21°44.9´w). norwegian hunting hut at the south end (botn = bottom) of loch fyne, built by the foldvik expedition in august 1926. it was also known as øiens hus and bunn huset. botten – see bundhytten i besselfjord. bottom terrace 73ø (73°24.4´n 23°15.0´w). name used by gunnar säve-söderbergh during the 1931–34 treårsekspeditionen for a terrace at the foot of stensiö bjerg, southern gauss halvø (sävesöderbergh 1933). (bottenterrassen.) boulder 71ø (71°37.2´n 25°16.1´w; map 5). prominent small nuna tak 3 km from the head of oxford gletscher, southern stauning alper. named by the 1970 university of dundee expedition which had a base camp on its top. the 1975 scottish expedition made use of the same site. boulder glacier 71ø (71°32.8´n 25°16.7´w). name occasionally used by the 1970 university of dundee expedition for oxford gletscher, south stauning alper; they established their base camp on a locality named boulder. uranus glacier has also been used. boulder ridge 74ø (74°19.9´n 24°36.4´w). ridge on the south side of djævlekløften, east clavering ø, where large boulders of per mian and crystalline rocks were found in a cretaceous sequence. the name was used by maync (1949). boulderbjerg 71ø (71°37.8´n 25°18.3´w; map 5). prominent peak about 2200 m high on the west side of upper oxford gletscher, south stauning alper. so named by the 1970 university of dundee expedition, which made the first ascent, because of its proximity to their base camp on boulder. (boulder peak.) bourbon ø 78ø-44 (78°45.6´n 18°13.5´w; maps 1, 5). island in the franske øer. named after kap bourbon by john haller during lauge koch’s 1956–58 expeditions. bow 72ø (72°12.2´n 24°43.3´w; map 5). mountain 1700 m high at the head of harlech gletscher, north stauning alper. first climbed by the 1963 imperial college expedition, and named after the london locality of bow, which originally had a bow-shaped bridge over the river lea. bowen bjerg 71ø-100 (71°41.9´n 22°05.3´w). mountain in east canning land. named by arne noe-nygaard during the 1931–34 treårsekspeditionen as mt. bowen after norman levi bowen [1887–1956]. a leading canadian petrologist and geochemist, he was noted particularly for his studies of assimilation in igneous magmas at the carnegie institute, washington d.c. boxøerne 78ø-16 (78°03.8´n 20°18.5´w; map 4). small islands in the southern part of jøkelbugten. so named by the 1906–08 danmark-ekspeditionen, probably because a depot of 37 boxes of dog-pemmican was made here during a snowstorm in february 1907. (boxöerne, box islands.) boyd bastion 73ø-703 (73°26.1´n 28°34.9´w). mountain forming the sw point of louise boyd land. named by john haller during lauge koch’s 1951 expedition. boykowdalen 72ø (72°55.5´n 22°26.9´w). valley on ne geogra phical society ø, equivalent to the present hundeklemmen. used only on nsiu maps (lacmann 1937), it was named after johann maria boykow [1879–1935], an austrian who gave instruction in photogrammetry, navigation and ballistic principles at the naval officers academy in berlin. brachiopoddal 74ø-148 (74°24.7´n 20°18.0´w). valley in western wollaston forland. so named by hans frebold during the 1931–34 treårsekspeditionen, for finds of fossil brachiopods. (brachio po dendal.) bragebreen 74ø (74°15.9´n 21°05.0´w). glacier on central clavering ø, a tributary to skillegletscher. used on nsiu maps (lacmann 1937), and named after brage, the poet-god of old nordic mythology noted for his wisdom. bramgåsesø 76ø-241 (76°49.6´n 19°02.9´w). small lake on winge kyst in southern germania land. named bramgaasesø by the 1906–08 danmark-ekspeditionen after the barnacle geese (= bram gås), which are common breeding birds in the region. bramgåssø 70ø-415 (70°29.6´n 27°56.6´w). small lake on sw milne land near rødefjord. named during the 1967–72 ggu scoresby sund expeditions by svend funder for the numerous young barnacle geese seen here. bramsen bjerg 74ø-125 (74°16.5´n 21°31.9´w). mountain ridge with three summits about 1270 m high on west clavering ø. named by lauge koch’s 1929–30 expeditions in the form mt. bramsen. the name has been applied to two different summits of the same mountain, but now covers the entire mountain. it is a common danish surname. (bramsens bjerg, bramsens bjærg.) brandaelv 73ø-168 (73°28.9´n 21°07.4´w). river on the south coast of hold with hope, named on an nsiu map (nsiu 1932a; fig. 13) as branda. there are many similar norwegian place names. brandal 72ø (72°48.8´n 22°31.2´w). valley on se geographical society ø, equivalent to the present adam af bremen dal. so named on nsiu maps of lacmann (1937), for the locality of the same name in sunnmøre, norway, home port of many sealers. brandalhytten 73ø (73°34.0´n 24°52.0´w). norwegian hunting hut in andrée land on the west side of geologfjord, built in september 1933 for arktisk næringsdrift; it has now disappeared. the hut was named after knut o. brandal, who helped build the hut, but died two weeks later of an acute illness; he was buried west of hoelsbu hunting station (p.s. mikkelsen 1994). it was also known as mørkebjerghytten. brandalvatnet 72ø (72°49.4´n 22°23.0´w). lake in adam af bremen dal (= brandal), geographical society ø. used on the nsiu maps of lacmann (1937). brandegga 72ø (72°51.0´n 22°27.8´w). mountain 726 m high on the north side of adam af bremen dal (= brandal), on se geo graphical society ø, equivalent to the present leitch bjerg. used on nsiu maps (lacmann 1937). bratskæret 76ø-178 (76°37.8´n 20°37.7´w; map 4). island in western dove bugt, perhaps rather large for a skerry. named by the eigil knuth’s 1938–39 mørkefjord expedition, for its appearance (brat = steep; skær = skerry). tutlas ø has also been used. brattegga 73ø (73°00.0´n 23°18.3´w). mountain 1260 m high on northern geographical society ø, south of rudbeck bjerg. so named on the nsiu maps of lacmann (1937), for the steep (= bratt) sides of the mountain. 141 bratthuken 72ø (72°28´n 21°59´w). the name is used in den grønlandske lods (1968), for the present takkerne, the ne point of ellemandsbjerge, eastern traill ø. it was named for the steep coastal cliffs. (kap bratthuken.) breccie elv 74ø (74°24´n 20°17´w). stream in brachiopoddal, western wollaston forland. the name was used in the geological report of rosenkrantz (1932) following work during lauge koch’s 1929 expedition, and was given for the brecciated nature of the rocks. (breccia river.) brede spærregletscher 80ø-52 (80°30.0´n 18°56.0´w; map 4). glacier on the south side of ingolf fjord, which extends across the fjord and partly blocks the sledge route to the interior. named by eigil nielsen during eigil knuth’s 1938–39 mørkefjord expedition as spærregletscher (brede = broad, spærre = block, obstruct). bredebræ 76ø-107 (76°42.0´n 22°48.0´w). broad glacier at the head of borgfjorden, formed by the confluence of storstrømmen and l. bistrup bræ; it produces many large icebergs. named as brede bræ by the 1906–08 danmark-ekspeditionen. (broad glacier, brede glacier, brei di jökull.) bredefirn 71ø-264 (71°58´n 24°11´w; map 5). glacier in the werner bjerge flowing west to join arcturus gletscher. named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk. bredefjord 75ø-36 (75°33.2´n 21°37.3´w; maps 2, 4; see also fig. 51). the broader of the two inner branches of ardencaple fjord, named brede fjord by the 1906–08 danmark-ekspeditionen. the narrower branch was named smallefjord. (brede fiord, breifjorden.) bredegletscher 70ø-76 (70°12.3´n 25°09.8´w; maps 3, 4). wide glacier on the south side of scoresby sund flowing north into vikingebugt. so named during the 1931–34 treårsekspeditionen by laurits bruhn. bredehorn 71ø-283 (71°53.3´n 23°59.7´w; map 5). mountain 1900 m high in the southern werner bjerge at the head of breithorn gletscher. it was originally named breithorn by peter bearth and eduard wenk during lauge koch’s 1953–54 expeditions (bearth 1959), after the swiss mountain of the same name. the name was initially approved as breidhorn, later becoming bredehorn. the mountain was climbed by bearth in 1954. see also breithorn gletscher. bredehorn gletscher – see breithorn gletscher. bredetop 74ø-402 (74°00.8´n 25°55.1´w; map 4). broad flattopped mountain in nw strindberg land, west of granitelv. named during lauge koch’s 1948–49 expeditions by hans r. katz for its broad summit. (breitstock.) bredevik – see breivik. bredgletscher 72ø-152 (72°20.1´n 22°43.0´w; map 4). glacier on se traill ø, south of mountnorris fjord. so named during lauge koch’s 1936–38 expeditions by hans peter schaub, because it is broad compared to nearby glaciers. bredningskærene 76ø-284 (76°22´n 21°18´w). group of skerries and islands on the east side of inderbredningen, east of rechnitzer land. so named by eigil knuth’s 1938–39 mørkefjord expedition because of their position in the expansion of a sound (= bredning). bredruphytten – see brædalhytten. bregnepynt 70ø-28 (70°55.5´n 25°23.6´w; map 4). peninsula on the ne coast of milne land, so named bregne pynt by carl ryder’s 1891–91 expedition because nikolaj hartz found numerous ferns (= bregne) here. brehytta 73ø (73°48.4´n 24°02.2´w). norwegian hunting hut in nord fjord, on the east side of waltershausen gletscher (bre = gletscher = glacier). it has also been known as solstrand, nordfjord and rød tophytten. brehytten – see bræhytten. breidhausen 73ø (73°26.0´n 22°05.0´w). mountain north of foldaelv in the giesecke bjerge, corresponding to the present bonney plateau. so named on an nsiu map (1932a), it derives from the norwegian for broad (= breid, brei). it was climbed in 1933 by noel odell and walter wood. breidifjòrdr 76ø (76°36.0´n 20°00.0´w). broad fjord, interpreted as possibly equivalent to the present dove bugt by tornøe (1944). the name is recorded in the icelandic sagas (bjørn jónssons grønlands annaler), and has been variously placed by early authorities. tornøe (1944) suggested this location in connection with his arguments for the site of another icelandic saga name, krosseyjar. breidvika – see breivika. breifjorden – see bredefjord. breithorn gletscher 71ø-157 (71°51.1´n 24°02.7´w; maps 4, 5). gla cier in the southern werner bjerge, draining into the nw end of pingo dal. the name was given by peter bearth and eduard wenk during lauge koch’s 1953–54 expeditions, and named after the mountain at the head of the glacier (breithorn, now bredehorn). it was approved in 1956 in the form bredehorn gletscher, the name replacing the rarely used biskop alfs gletscher. however, in 1971 the name was officially changed to breithorn gletscher to conform with the common use of this form in scientific publications, although the mountain bredehorn has retained its danicised name. breivik 74ø-252 (74°05.9´n 21°07.0´w). norwegian hunting hut on the south side of clavering ø, east of eskimovig. originally built on west clavering ø by the foldvik expedition in 1927, it was moved to this site in the summer of 1929. it was named after the bay which the norwe gians called breivika. a newer norwegian hut on the same site, known as breivikhytten, was probably built in august 1938. breivika 74ø (74°05.7´n 21°07.5´w). bay on south clavering ø, equivalent to the present eskimovig. so named in this form on 1932 nsiu maps, and as breidvika on the later maps of lacmann (1937). named for the form of the bay (breid = broad, wide). breivikdalen 74ø-251 (74°06.4´n 21°07.5´w). valley on south clav ering ø. so named on the nsiu (1932a) map, after the hunting hut (breivik) and bay (breivika) at the mouth of the valley. on lacmann’s (1937) maps breidvidalen is used. (breidvikadalen.) bremsholmane 72ø (72°44.3´n 21°49.3´w). line of skerries off se geo graphical society ø. the skerries form a hinderance or brake (= brem se) to the winter-ice. so named on the nsiu maps of lacmann (1937). brescia hill 70ø (70°04.8´n 23°06.2´w). name used by leonardo bonzi’s 1934 expedition during their exploration of volquaart boon kyst, most probably for the summit west of their ghiaccaio brescia. breslauer spids 71ø (71°53.4´n 25°35.0´w; map 5). mountain about 2510 m high between hecate gletscher and the upper part of spærre gletscher, southern nathorst land. climbed by karl m. herligkoffer’s 1966 expedition on 23 august, and named after the town of breslau/wroclaw in sw poland. the peaks on this mountain ridge are also known as silberspitzen. brillen 76ø (76°44.2´n 20°43.7´w). two islands sw of vindseløen are so named on c.s. poulsen’s (1991) map in his published diary of the 1906–08 danmark-ekspeditionen (j. løve, personal communication 2009). they may correspond to the present ringøen and midter holmen. brinkley bjerg 74ø-11 (74°09.5´n 20°45.5´w). mountain 1075 m high on se clavering ø, named by william scoresby jr. in 1822 as cape brinkley. it probably commemorates john brinkley [1763–1835], bishop of cloyne, first astronomer royal for ireland, and professor of astronomy at dublin. scoresby’s cape was probably the mountain to which the name was transferred by the place name committee in about 1935. brinkley plateau 74ø-230 (74°08.9´n 20°45.4´w). plateau on se clavering ø from which brinkley bjerg rises. first used by lauge koch’s 1929–30 expedition, originally in the form mt brinkley 142 plateau. brinkmann fjæld 70ø (70°39.5´n 22°42.8´w). minor summit on the west side of hurry inlet, north of muskusoksekløft. the name was used by aldinger (1935) in his report on work during the 1931–34 treårseks peditionen, and was given for r. brinkmann, a german geologist with interests in jurassic ammonites. (brinkmannfjæld, brinkmann fjaeld.) brisbane bjerg 74ø-10 (74°12.5´n 20°09.6´w). mountain 486 m high on east clavering ø, named by william scoresby jr. in 1822 as cape brisbane in compliment to sir thomas makdougall brisbane [1773–1860]. a noted astronomer, brisbane was president of the royal society of edinburgh. scoresby’s cape was later found to be a mountain, and the name changed accordingly. bristol elv 72ø-234 (72°27.5´n 22°30.9´w). river on eastern traill ø, draining south into mountnorris fjord. so named by desmond donovan during lauge koch’s 1949–50 expeditions, after the town of bristol in england. donovan was at bristol university. britannia gletscher 77ø-121 (77°11.0´n 24°00.0´w; maps 2, 4; fig. 21). large glacier in north dronning louise land flowing into britannia sø. the name is derived from the roman name for an cient britain, and was given by the 1952–54 british north green land expedition for patriotic reasons; the expedition was british and queen elizabeth ii had recently succeeded to the throne. the name unicorn gletscher was apparently used during the expedition, and is still occasionally encountered in correspondance (e.g. p.s. mikkelsen 1994). britannia sø 77ø-100 (77°08.6´n 23°24.6´w; map 4). large lake in northern dronning louise land, named by the 1952–54 british north greenland expedition. the expedition had its main base on the north shore of the lake, but the buildings were destroyed by the advance of britannia gletscher in the 1980s. in 1951 admiralty lake was used, and the name slamsøen is occasionally encountered. britta dal 73ø-114 (73°25.1´n 23°16.9´w). ravine on the sw coast of gauss halvø between stensiö bjerg and smith woodward bjerg. so named by gunnar säve-söderbergh during the 1931–34 treårsekspeditionen, after his wife britta kerstin arnell [d. 1952]. broad ridge 73ø (73°31.0´n 23°22.3´w). locality between two ravines on the south side of sederholm bjerg, gauss halvø. the name was used in a report on work during the 1931–34 treårseks peditionen (johansson 1935). (breda ryggen.) brocchieridalen 72ø (72°57.0´n 24°32.0´w). valley on western geographical society ø. so named on the nsiu maps of lacmann (1937) for vittorio beonio-brocchieri [b. 1902], a professor at the university of pavia who took part in the 1930 nsiu expedition. broch øer 72ø-68 73ø-273 (73°04.0´n 22°34.0´w; maps 3, 4). group of islands off ne geographical society ø and the mouth of sofia sund, including langåren, borgøen, tveholmen, skild pad den, søstjernen, kamelen and other unnamed islands. the name broch inseln had originally been given by karl koldewey’s 1869–70 expedition to the present vinterøer, although the name is only found on the maps in payer’s (1876) narrative. the 1899 a.g. nat horst expedition (nathorst 1901) transferred the name to the present site, but this was probably unintentional. the islands were probably named after ole jacob broch [1818–1889], a norwegian mathematician, physicist and politician. (brochs öar, broch islands, brochöyane, brocks øer.) brockmeyer bjerg 80ø-20 (80°27.5´n 21°27.8´w). mountain in kronprins christian land. the name was given to a distant nuna tak by the 1909–12 alabama expedition, and commemorates commander brockmeyer of the danish inspection ship islands falk that had assisted the expedition on several occasions. ernst jens gustav brockmeyer [1862–1940] had provided ejnar mik kelsen with an engineer, iver p. iversen, to replace his own engineer (j. løve, personal communication 2009). the place name com mittee could not locate the mountain with certainty, but wishing to retain the name placed it in an area which then had relatively few names; the only map in their archives suggesting a location places it just south of sødalen. on modern maps the 1028 m ice-capped summit south of sødalen and ne of keglen is the most appropriate location. (brockmeyer’s nunatak.) broer ruys nord 73ø (73°32.7´n 20°29.7´w). danish hunting hut at the mouth of glommen, about 4 km ne of kap broer ruys, built by nanok in september 1945. it has often been known as dom kirken. it is close to, and slightly south of the norwegian hut known as skandalen or bukta. broer ruys station – see kap broer ruys station. broer ruys syd 73ø (73°27.5´n 20°53.7´w). danish hunting hut on the south coast of hold with hope, sw of kap broer ruys. it was built by nanok in 1945. brogetdalen 73ø-598 (73°45.8´n 24°48.8´w; map 4). large valley in strindberg land draining east to nordfjord. the name was first used by teichert (1933) during the 1931–34 treårsekspeditionen, and is a translation of his original buntes tal (= painted valley) given for the extravagant colours of the rocks (broget = multicoloured). the map of giæver (1939) indicates a hunting hut in brogetdalen about 10–15 km inland, but this was never built (p.s. mikkelsen 1994). the names strindberg valley, stordalen and giæverdalen have also been used for the valley. (brogede dal.) brorson halvø 74ø-342 (74°37.0´n 19°28.0´w). northern peninsula of wollaston forland. the name appears to have been given by the place name committee in 1939, possibly after hans adolf brorson [1694–1764], a danish bishop and noted hymn writer. maync (1947) used the form brorsons halbinsel. brorson pynt 76ø-89 (76°41.6´n 18°32.4´w). south cape of måt ten, an island south of kap bismarck. named by the 1906–08 danmark-ekspeditionen as brorsons pynt, perhaps for h.a. bror son. see also brorson halvø. brown-stua 71ø (71°45.9´n 22°31.8´w). norwegian hunting hut built by the møre expedition in fleming fjord in august 1931, about 5 km sw of kap brown. it is also known as holstad. (kap brown hytten, brownhuset.) broxdalen 72ø (72°56.2´n 24°13.7´w). valley on western geo graphical society ø. so named on the nsiu maps of lacmann (1937) after the norwegian telegraphist leif brox [b. 1905], who was stationed at myggbukta from 1928 to 1930. bruce fjord 69ø-71 (69°52.0´n 23°04.0´w; map 4). fjord on the north side of manby halvø, northern blosseville kyst. named by malcolm slesser’s 1969 expedition for william speirs bruce [1867–1921], a polar explorer and oceanographer, who took part in expeditions to the antarctic in 1902–04 and spitsbergen in 1906–20. bruddal 70ø-290 (70°30.0´n 22°13.2´w). valley in south liverpool land, so named by alfred rosenkrantz during lauge koch’s 1926– 27 expeditions because the valley follows a fault zone (= brud). (fault valley.) brudelv [kuukajik kittikajik] 70ø-296 (70°30.0´n 22°13.2´w). river in south liverpool land draining bruddal. named during the 1931–34 treårsekspeditionen by laurits bruhn. brumse gletscher 72ø-154 (72°19.0´n 22°34.1´w). glacier on se traill ø, south of mountnorris fjord. named during lauge koch’s 1936–38 expeditions by hans peter schaub for the lead dog in his sledge team. (brumsegletscher, brumses gletscher.) brun bjerg 73ø-315 (73°51.0´n 23°10.0´w). mountain in central hudson land, north of ritomsø. named by heinrich bütler during lauge koch’s 1936–38 expeditions for albert brun, a naturalist who visited spitsbergen in the early 1900s. (brunberg, bruns bjerg.) brune nunatakker 71ø-428 (71°09.1´n 29°37.4´w; map 4). nuna tak group west of graben land. so named by peter home wood during the 1967–72 ggu scoresby sund expeditions be cause of their characteristic brown (= brune) colour. 143 brunedal 70ø-413 (70°40.5´n 28°06.0´w; map 4). valley on the west side of rødefjord. named during the 1967–72 ggu scoresby sund expeditions by kai sørensen, for the dominant brown-weathering colour of the rocks. brünhildbreen 74ø (74°25.2´n 21°09.4´w). small glacier on north clavering ø. used on the nsiu maps of lacmann (1937), the name commemorates brunhild (brunhilde), the queen of outstanding strength and beauty who married gunther in the the german epic poem from c. 1200, the nibelungenlied. brunhild also features in old norse literature. brünhorn 73ø (73°32.5´n 25°56.0´w). peak 2066 m high on the south side of grejsdalen, andrée land. climbed by the 2007 army boreal zenith expedition. brunknøs 75ø-76 (75°16.7´n 22°07.4´w). mountain in c.h. ost en feld land. the name originated from the wintering party at kulhus during the 1931–34 treårsekspeditionen, and was given for the abundant occurrence of rusty brown rocks. brygger elven 75ø (75°08.8´n 19°49.3´w). name used by danish hunters in the 1930s for a minor river west of nanok hunting station, hochstetter forland (nyholm-poulsen 1985). bryozoen-ryg 73ø (73°33.8´n 22°03.0´w). ridge in ne giesecke bjerge where bendaelv begins. the name was used by maync (1942), and was given for finds of fossil bryozoa. brystet 76ø-53 (76°57.4´n 20°18.4´w). conical hill 412 m high at the se end of sælsøen, so named by the 1906–08 danmarkekspeditionen because of its shape. (the breast.) brystværnene 73ø-548 (73°02.0´n 28°01.7´w). mountain ridge about 2000 m high in northern goodenough land, east of nordenskiöld gletscher. it was named by j.m. wordie’s 1929 expedition as battlements for its appearance. brædal 75ø-81 (75°34.0´n 21°26.8´w). valley in sw dronning margrethe ii land draining into bredefjord. the name originated from the wintering party at kulhus during the 1931–34 treårs ekspeditionen, and was given for the many glaciers which drain into the valley (bræ = glacier). brædalhytten 75ø (75°32.8´n 21°28.1´w). norwegian hunting hut east of the mouth of brædal, on the north side of bredefjord, built in august 1933 by john giæver’s expedition. it was also known as bredruphytten and solstrand, but had completely disappeared by 1988. bræfjorden 76ø-166 (76°25.3´n 22°00.0´w; map 4). fjord on the north side of rechnitzer land. the name was reported as used by nanok hunters, and appears first on a map in jennov (1939). it was named for the two picturesque glaciers on the south side of the fjord. (bræfjord.) bræfjordhytten 76ø-204 (76°28.5´n 21°41.2´w). danish hunting hut on the north side of the mouth of bræfjorden, built by nanok in may 1934. now a ruin. it has occasionally been referred to as jarnershytte. bræhytten 73ø (73°09.6´n 27°33.8´w). norwegian hut east of the mouth of knækdalen, south frænkel land, built in april 1950. the inner part of the fjord is often blocked by ice calved from nordenskiöld gletscher (bræ = glacier). the hut has also been known as gregorydalhytten and knækelvhytten. (brehytten, bræhytten.) brændestablen 70ø-20 (70°27.4´n 28°06.1´w). small skerry just off the south point of rødeø in rødefjord, so named by carl ryder’s 1891–92 expedition. a prominent dolerite dyke here is divided by joints into horizontal prismatic blocks which resemble a pile of firewood (= brændestablen). bræøerne 76ø-114 (76°45.4´n 22°10.2´w; map 4). three islands off the front of bredebræ at the west end of borgfjorden. so named by the 1906–08 danmark-ekspeditionen. the southernmost island has been referred to as southern bræ ø. (bræöerne, glacier islands, bræ-inseln, bræöerne, jökuleyjar.) brøggerhytta 73ø (73°15.4´n 23°59.3´w). norwegian hunting hut on the north side of dusén fjord, gunnar andersson land, east of zoologdalen. built by arktisk næringsdrift in september 1929, the name commemorates a lawyer named brøgger (p.s. mikkelsen 1994). (brødger, brøgger-hytta, brœgger-hytta, brøggers hytte.) brønlunds grav 79ø-4 (79°09.3´n 19°04.0´w; maps 1, 4). point on the east coast of lambert land, so named by the 1906–08 dan mark-ekspeditionen because jørgen brønlund died here about november–december 1907, one of the three expedition members to die on their return from their northern explorations. jørgen brønland [1877–1907], a greenlandic member of the expedition, had previously taken part in the 1902–04 literary expedition to north-west greenland. a relief party found his body in march 1908. in 1909 ejnar mikkelsen also visited the site. eigil knuth (1940) records that lambert land rises behind the grave like a mighty burial mound. the site was visited in april 1963 by members of the sirius sledge patrol, who erected a bronze memo rial plaque, a gift from knud lauritzen, and re-buried the remains of the body under a large cairn. on the 70th anniversary of the 1906–08 danmark-ekspeditionen in 1978 a group of four led by jørgen bjerre erected another memorial plaque. (brönlunds grav.) braastad 73ø (73°36.5´n 22°28.5´w). norwegian hunting hut on the north side of moskusoksefjord, east of ankerbjerg. built in september 1929 by arktisk næringsdrift, the hut was named after johan braastad [b. 1888], geologist and secretary of nsiu from 1924–1935. it is also known as ankerlien. (bråstad, bråstedhytte.) buache ås 71ø-23 (71°54.6´n 22°47.3´w). ridge in eastern scores by land north of fleming fjord. it was named by william scoresby jr. in 1822 as cape buache in compliment to a french philosopher, probably jean nicholas buache [1741–1825]. buch bjerg 71ø-25 (71°31.7´n 22°34.1´w; map 4). mountain 770 m high on the west side of carlsberg fjord. named by william scoresby jr. in 1822 as cape buch after the celebrated geological traveller baron christian leopold von buch [1774–1853]. buch was considered to be the most illustrious geologist that germany produced in the 19th century. scoresby’s cape was evidently a mountain and the name was later changed accordingly. buddha 71ø-429 (71°12.2´n 28°09.0´w; map 4). mountain 1880 m high ne of graben land. so named by johan d. friderichsen during the 1967–72 ggu scoresby sund expeditions because from the sw it resembles a buddha. buddingbjerg 73ø-407 (73°19.8´n 25°54.9´w). mountain 1805 m high between benjamin dal and junctiondal, southern andrée land. named during the 1948–50 lauge koch expeditions by erd hart fränkl for its rounded shape and layered appearance, resembling a pudding (= budding). it was climbed by fränkl and fritz schwarzenbach in august 1950. budolfi isstrøm 76ø-131 (76°19.0´n 25°00.0´w; maps 2, 4; fig. 21). large glacier in carlbergfondet land, southern dronning louise land, flowing eastwards to join l. bistrup bræ. named by j.p. koch during his 1912–13 expedition, perhaps for saint bu dolfi, patron of a church in ålborg, denmark. the 1952–54 british north greenland expedition that traversed the glacier considered the term ‘isstrøm’ particularly unsuitable for this glacier, because it moves only slowly. a.b. drachmann glet scher has also been used, but this name is now applied to a more southerly glacier. (budolfi gletscher, budolfs skriðjökull.) buegletscher 72ø-158 (72°17.8´n 22°31.0´w). glacier on se traill ø, south of mountnorris fjord. named during lauge koch’s 1936–38 expeditions by hans peter schaub for its curved shape (bue = bow, curve). bukta 73ø (73°33.3´n 20°30.5´w). norwegian hunting hut in a bay (= bukta) on the east coast of hold with hope, nw of kap broer ruys, built by the foldvik expedition in august 1927. it was also known as skandalen and moskusoksehytten. bulbjerg 70ø-275 (70°03.8´n 22°51.5´w). mountain on volquaart boon kyst. so named by laurits bruhn during the 1931–34 tre 144 årsekspeditionen, after the prominent cliff of the same name in nw jylland, denmark. it was climbed by the 1934 bonzi expedition and named punta roma. bültrop fjelde 77ø-61 (77°25.8´n 20°32.3´w; map 4). mountains on the south side of v. clausen fjord, inner skærfjorden. so named by david malmquist during the 1931–34 treårsekspeditionen, after the husband of his wife’s sister, the mathematician einar bül trop lunell. he was professor at the university of umeå, sweden. bundermannfjellet 74ø (74°21.0´n 20°47.7´w). mountain 1369 m high on west clavering ø, corresponding to koralbjerg. so named on the nsiu maps of lacmann (1937) for max bundermann [b. 1904], who took part in photogrammetric work on nsiu aerial photographs of east greenland. bundfjeldet 76ø-230 (76°58.3´n 21°48.2´w). innermost mountain on the north side of vigfus dal, daniel bruun land. named by eigil knuth’s 1938–39 expedition for its position at the end of the fjord (bund = bottom, also the inner part of a fjord or bay). bundhuset 71ø (71°37.8´n 22°59.8´w). norwegian hunting hut built in august 1932 by helge ingstad’s expedition about 3 km from the inner end of fleming fjord. it has also been known as heimen and ingstadheimen. bundhuset – see bunnhuset. bundhytten – see inderhytten. bundhytten 75ø (75°20.1´n 20°11.8´w). danish hunting hut on the north (inner) side of peters bugt, north of the mouth of arden caple fjord, built for nanok in august 1930. it is officially known as petersbugthytten, and has also been called nummer 1 hytten. bundhytten i besselfjorden 75ø (75°59.2´n 21°53.3´w). norwegian hunting hut built by john giæver’s expedition in innermost bessel fjord in august 1932. it was a ruin in 1989. (botten, bundhytten.) bundhytten i tyrolerfjorden 74ø (74°36.6´n 22°05.4´w). norwegian hunting hut built in september 1932 by sigurd tolløfsen’s expedition about 1 km inland from the innermost part of tyrolerfjord, northern payer land. (tyrolerfjord bundhytte, fjordbotten). bundstykket 75ø (75°59.5n 21°37.9´w). peninsula and mountain on the north side of inner bessel fjord, that charles poulsen com pared to danmarksmonumentet in mørkefjord during the 1906– 08 danmark-ekspeditionen (poulsen 1991). bunn-huset 71ø (71°38.0´n 22°23.7´w). norwegian hunting hut built in august 1932 by helge ingstad’s expedition in the inner part of nathorst fjord. it has also been referred to as siste-huset. bunnhuset 73ø (73°40.6´n 21°44.9´w). norwegian hunting hut at the south end of loch fyne, built in august 1926 by the foldvik expedition, and also known as øiens hus. (botnhuset, bundhuset, bundhytten.) bunnhuset 73ø (73°19.1´n 25°02.8´w). hut at the west end of dusén fjord, ymer ø, built in august 1932 by the crew of the isbjørn for salmon fishing, and subsequently also used by hunters. it has also been known as noahytten, laksehytten and holmboe-hytta. buntes tal 73ø (73°45.8´n 24°48.8´w). original name for broget dalen in strindberg land, given by curt teichert in 1931 because the colour effects of the rocks in the steep walls of the valley were reminiscent of those he had seen in the painted desert of colorado and utah. teichert considered the official name brogetdalen (= the multicoloured valley) did not adequately convey the extravagance of colour. buri søer 72ø-461 (72°41.2´n 27°39.9´w). group of lakes in niklausdal, western gletscherland. named during the 1931–34 treårsekspeditionen by eugène wegmann, for a geologist of this name at zürich, who subsequently became professor. the association with niklausdal is said to be significant, a ‘klaus’ in swiss dialect being a simple character (fritz schwarzenbach, personal communication 1996). buskøysundet 72ø (72°46.9´n 22°55.5´w). sound between gåseøen and kista ø in vega sund. used on the nsiu maps of lacmann (1937), the name commemorates the buskø, a norwegian sealer used by arktisk næringsdrift expeditions to east greenland. (busk öy sunde.) bütlers klippe 72ø (72°09.5´n 23°45.7´w). name used on preliminary map sheets of the mesters vig region, for a cliff about 100 m above sea level; it was changed on the published maps printed in 1951 to the present permklippen (e.g bondam 1955). the name was given by prospecting teams associated with lauge koch’s 1948–49 expeditions after heinrich bütler, a swiss geologist who worked for many years in east greenland with lauge koch’s expeditions. bælgen 76ø-15 (76°20.1´n 20°14.8´w; map 4). ne cape of nanok ø, so named by the 1906–08 danmark-ekspeditionen. achton friis and aage bertelsen camped here for 14 days, and the name may derive from the windy and exposed location.the island nanok ø has a bellows-like shape on a map (j. løve, personal communication 2009). (kap bælgen, bælget, the bellows.) bæltenunatak 70ø-446 (70°11.2´n 29°47.0´w). nunatak on the se side of vestfjord gletscher. so named by w.e. adrian phillips during the 1967–72 ggu scoresby sund expeditions because it is cut by a n–s-trending belt of black rocks (bælte = belt). bændelbjerg 70ø-449 (70°25.4´n 29°46.6´w). mountain 2341 m high in west paul stern land. named by w.e. adrian phillips during the 1967–72 ggu scoresby sund expeditions for the tapelike striped rocks on its western side (bændel = tape). bærtun 72ø (72°26.7´n 25°28.9´w). norwegian hunting hut built by the møre expedition in september 1931 on the north side of fors blad fjord, west of polhem dal. it was named for the berries (= blok kebær; rogne 1981). the hut has also been known as pol hems dalhytten. bødtker-hytta 73ø (c. 73°01´n 23°38´w). norwegian hunting hut about 10 km west of rudbeck bjerg, northern geographical society ø, a locality known to norwegians as kapp veslekari. the hut was built here in september 1929 by arktisk næringsdrift, moved to the opposite side of the fjord in 1930 where it was known as stor-dalen, and moved again in 1931 to renbugten where it was called reinsbukta. (bødker, sejerstedt bødtkers hytte.) bøggild bjerg 73ø-78 (73°29.1´n 22°56.4´w). mountain on gauss halvø. named by lauge koch’s 1929–30 expeditions in the form mt. böggild after ove balthazar bøggild [1872–1956], a danish geologist and mineralogist, noted particularly for his studies of cryolite. (bøggilds bjerg.) bøllebakken 74ø (74°28.1´n 20°31.8´w). feature se of zackenberg forskningsstation. the name has been used by visiting scientists. börge elv 77ø-77 (77°32.5´n 19°12.2´w; map 4). river draining the se part of stormlandet. named during the 1931–34 treårseks peditionen by david malmquist in the form börje elv after börje uhnno, a friend of malmquist who became a medical superintendant at gävle hospital, sweden. børsnæsset 80ø (80°45.6´n 14°15.0´w). cape on the east coast of amdrup land where the sledge parties of the 1906–08 danmarkekspeditionen split up. their provisions were divided and the remainder placed in a depot (børs = stock exchange; j. løve, per sonal communication 2009). børøya 72ø (72°42.3´n 22°46.3´w; fig. 14). island in central vega sund, the present silja ø. so named on the nsiu maps of lacmann (1937) for an island of the same name in vesterålen, norway. (böröya.) bøygen 72ø (72°00.0´n 24°59.3´w; map 5). summit about 2200 m high between col des pulkas and grantalang col, stauning alper. climbed by the 1996 norwegian stauning alper expedition, and named after the supernatural being in ‘peer gynt’ by henrik ibsen. bådhytten 72ø (72°24.5´n 23°34.6´w). hut built on the south coast of traill ø in the summer of 1968 by personnel from mestersvig airfield. it was constructed from an old boat, the polypen, formerly owned by lauge koch’s expeditions and based at koch’s research station on ella ø. 145 bådskæret 76ø-68 (76°45.5´n 18°47.6´w). small island or skerry off wendel pynt, west of danmark havn. named by the 1906–08 danmark-ekspeditionen as baadskæret, apparently because of inuit stone ruins found here interpreted as supports for kayaks (båd = boat). according to friis (1909) the skerry was initially called hvalrosskærene. baadskær 77ø (77°16.9´n 18°20.1´w). name used by c.s. poulsen during the 1906–08 danmark-ekspeditionen for a skerry off eastern rosio, ne germania land (poulsen 1991). the boat from the first boat trip was laid up here because further progress was blocked by ice (j. løve, personal communication 2009). bådsted 74ø-185 (74°05.8´n 21°02.8´w). small bay east of eskimo vig, south clavering ø. the name was used as a botanical reference locality in reports of the 1931–34 treårsekspeditionen in the form baadsted (gelting 1934); it was said to be a good harbour for small boats. c c. drost ø 77ø-28 (77°36.8´n 20°31.0´w; map 4). island at the inner end of penthievre fjord. so named by the 1906–08 danmarkekspeditionen, probably for carl drost [1854–1926], a businessman and ship-owner. (c. drosts ö.) c.f. knox tinde 72ø-509 (72°05.2´n 24°51.8´w; map 4). moun tain about 2750 m high at the head of bersærkerbræ, gully glet scher and schuchert gletscher. first climbed by a cambridge uni versity expedition on 22 july 1963, it is best known in mountaineering literature under the name grandes jorasses, the name originally proposed by malcolm slesser following his 1958 expedition. the name was changed in november 1964 to commemorate colin frederick knox [1938–64], a new zealand climber who led the 1963 cambridge expedition, and who died the following year in the french alps. the second ascent was made by an imperial college expedition in august 1963. (knoxtinde.) c.f. mourier fjord 77ø-34 (77°21.0´n 20°16.7´w; map 4). fjord in the sw part of skærfjorden. named by the 1906–08 danmarkekspeditionen, probably after christian frederik denys mourier [1879–1957], a lieutenant in the danish navy (j. løve, personal communication 2009). c.h. jørgensen nunatak 80ø (c. 80°40´n 22°20´w). mountain in kronprins christian land. named by the 1909–12 alabama expedition after christian h. jørgensen, a lieutenant in the danish army and one of the expedition members. initially approved, this name was subsequently discarded because of the difficulty of identifying the original feature. c.h. ostenfeld land 75ø-93 (75°14.0´n 21°30.0´w; maps 2, 4). land area between grandjean fjord and ardencaple fjord. map ped in part by lauge koch during flights in 1932 on the 1931–34 treårsekspeditionen, it was named after christian emil hansen ostenfeld [1873–1931], a danish botanist noted for his ‘flora of greenland and its origin’. ostenfeld was chairman of the carlsberg foundation that supported the 1931–34 treårsekspeditionen. (c.h. ostenfelds land.) c.h. ostenfeld nunatak 74ø-142 (74°17.2´n 22°55.6´w; map 4). large nunatak in wordie gletscher, named by lauge koch’s 1929– 30 expeditions. see also c.h. ostenfeld land. (c.h. ostenfelds nunatak.) c. hoffman halvø 70ø-400 (70°57.0´n 27°45.0´w; map 4). penin sula between harefjord and rypefjord. named by the 1963 geo dætisk institut expedition after the helicopter mechanic, c. hoff man, who was killed here when he walked into a rotor blade. c.j. ring fjelde 80ø-115 (80°15.0´n 18°55.5´w; map 4). mountain range on the north side of hekla sund. named by john haller following explorations during lauge koch’s 1956–58 expeditions. carl johan ring [1870–1918] was the norwegian ice-pilot on the 1906–08 danmark-ekspeditionen, and had previously sailed on the expedition ship as 1st mate when it went under the name the magdalena. as an experienced skier he took part in many of the most demanding depot-laying journeys during the 1906–08 dan mark-ekspeditionen. c. mountain 72ø (72°48.0´n 27°27.1´w). mountain in gletscher land, the present lugano bjerg. this was a temporary designation used by louise boyd’s 1931 expedition (boyd 1935). c. silfverberg ø 77ø-29 (77°34.0´n 20°07.7´w; map 4). island be tween penthievre fjord and agsutsund. named by the 1906–08 dan mark-ekspeditionen as c. silfverbergs ö, possibly for conrad emil silfverberg [1875–1941], a lieutenant in the danish navy, who from 1902 worked for a salvage company. (silferbergs ø, silverbergs ø.) cai torino 72ø (72°12.0´n 25°07.2´w). peak about 2000 m high in the northern stauning alper west of frihedsgletscher. it was climbed by g. dionisi’s 1982 expedition, and named after the turin branch of the italian alpine club (cai = club alpine italiano). cadegnosø 73ø-310 (73°51.0´n 23°15.5´w). lake in central hud son land. named by heinrich bütler during lauge koch’s 1936–38 expeditions for the cadegnosee, a lake in the st. gotthard region of switzerland. caerleon 72ø (72°14.3´n 24°37.8´w). mountain 2028 m high on the north side of bersærkerbræ, north stauning alper. named by john hunt’s 1960 expedition after caerleon castle, wales, a roman legionary fortress and site of a norman castle associated with the legends of king arthur. hunt’s party abandoned their ascent close to the summit when a cornice gave way. the 1963 imperial college expedition claimed the first ascent. bennet (1972) noted caerleon as identical with tårnfjeld, while slesser (1964a, b) considered them to be different summits, although close to each other. caerleon glacier 72ø (72°12.6´n 24°35.7´w). glacier on the north side of bersærkerbræ, north stauning alper, equivalent to the present tårnfjeld gletscher, with tårnfjeld (caerleon) at its head. so named by john hunt’s 1960 expedition. caius fjeld 72ø-505 (72°05.3´n 25°11.3´w; map 5). caius fjeld and gonville fjeld are two sharp rock summits each about 2280 m high on the west side of cavendish gletscher, stauning alper. first climbed by the 1963 cambridge university expedition, this peak was named after caius college, cambridge (properly gonville and caius), founded by edmond gonville in 1348 and refounded by dr. caius in 1557. calamites dal 71ø (71°44.2´n 22°30.6´w). valley on the se side of wegener halvø in which calamiteselv flows. the name is used occasionally in geology reports. calamiteselv 71ø-53 (71°44.21´n 22°30.6´w). river on the se side of wegener halvø, named by lauge koch’s 1926–27 expeditions as calamites river for finds of fossils. calamiteselv 72ø-209 (72°11.8´n 23°49.3´w; map 5). river draining north from lille blydal into noret, west of mesters vig. named by prospecting teams associated with lauge koch’s 1948–49 expeditions, for the fossil finds. calcitdalen 70ø-429 (70°21.1´n 26°49.6´w; map 4). east–westtrending valley in eastern gåseland, draining into gåsefjord. named during the 1967–72 ggu scoresby sund expeditions by georg sawatzki for several occurrences of limestone. caledoniahytten 72ø (c. 72°25´n 25°53´w). norwegian hut said to have been built in 1931 south of caledoniaø in forsblad fjord. it has also been referred to as sulebak. according to p.s. mikkelsen (1994) the hut was never built. caledoniaø 72ø-131 (72°25.0´n 25°48.6´w). island in forsblad fjord, named by helge g. backlund in 1929 at the suggestion of his assistant (arne noe-nyegaard) as caledonia island. the island lies in an area influenced by orogenic (mountain building) activity of caledonian age. cambridge bugt 72ø-73 (72°48.5´n 22°00.0´w; maps 3, 4). large bay on the east side of geographical society ø. named by j.m. 146 wordie’s 1926 expedition as cambridge bay. all three of wordie’s east greenland expeditions, in 1923, 1926 and 1929, were sent out under the auspices of the university of cambridge. cambridge toppe 73ø-533 (73°04.1´n 27°42.6´w). series of summits about 2200 m high in northern goodenough land, named by j.m. wordie’s 1926 cambridge expedition as cambridge peaks. the peaks were climbed in 1926, and on the maps of his 1929 cambridge expedition (wordie 1930a, b) five summits on the ridge were included under this general name. camp col 71ø (71°37.8´n 25°20.5´w; map 5). pass at the head of oxford gletscher leading over to triton glacier. camp creek 73ø-288 (73°19.2´n 22°43.6´w). stream draining the south coast of gauss halvø east of koralkløft. so named by gun nar säve-söderbergh during the 1931–34 treårseks pedi tionen. camp lindquist 72ø (72°53.6´n 24°47.3´w). norwegian hunting station built in 1930 by arktisk næringsdrift on ne ella ø, about 3 km south of kap elisabeth. gustav lindquist helped to build the station, and over-wintered here 1930–31. it has also been known as maristua. camp river 70ø (70°26.8´n 23°00.0´w). minor river in south jame son land flowing into hesteelv. so named by hermann aldinger during the 1931–34 treårsekspeditionen, because his camp was sited by the river (aldinger 1935). camp peninsula 77ø (77°36.1´n 20°42.0´w). peninsula on the south side of klægbugt, nordmarken, the site of the base camp of the 1987 irish expedition to northern east greenland. camp tahoe 72ø (72°13.3´n 24°03.2´w). name used by washburn (1965) for the house north of tunnelelv on a section of road between nyhavn and minebyen, the headquarters for his geomorphological studies between 1955 and 1964. it was named after his wife, tahoe washburn. it is usually known as washburns hus. campanulavig 77ø-72 (77°40.0´n 20°31.0´w). small inlet or fjord in the inner part of skærfjorden. so named by david malmquist during the 1931–34 treårsekspeditionen, because of the finds of ‘campanula uniflora’ here, which were then the northern limit of the species in east greenland. the name has usually appeared on maps in the form campanulavigen. campbell sund 71ø-5 (71°17.5´n 21°46.1´w; map 4). sound between trekanten and the mainland on the coast of north liver pool land, originally named campbell bay by william scoresby jr. in 1822. this name, and others in the vicinity, were given by scoresby for friends chiefly resident in manchester. terminal sund has also been used. (campbell sound, campbell bucht.) campo venezia 72ø (c. 72°12´n 24°50´w). name used by braun (1953) for a camp site on skjoldungebræ, north stauning alper, from which ascents of four peaks were made during lauge koch’s 1951 expedition. the camp was surrounded by a system of glacial streams, compared to the canals of venice. canis major gletscher 71ø (71°41.1´n 25°23.0´w; map 5). upper tributary of jupiter gletscher, southern stauning alper, so named by the 1975 scottish scoresby land expedition for the constellation. see also kap canis major. canis minor gletscher 71ø (71°39.5´n 25°17.7´w; map 5). upper tributary of jupiter gletscher, south stauning alper, so named by the 1975 scottish scoresby land expedition for the constellation. see also kap canis minor. canning land 71ø-12 (71°41.5´n 22°12.0´w; maps 3, 4; see also fig. 90). major peninsula between nathorst fjord and carlsberg fjord. william scoresby in 1822 had used the name canning island for a tract of bold land that appeared to be insular (fig. 3). it was named after one of the secretaries of state, george canning [1770–1827], member of the british parliament for liverpool from 1812–1822. hartz (1902) recorded that on sailing north in the antarctic on 23 august 1900 canning island re vealed itself as a peninsula and was renamed canning land. (canning ø, cannings ö, canning peninsula, canningland, caning land). canongletscher 75ø-75 (75°28.4´n 22°27.5´w; map 4; see also fig. 81). glacier sw of the head of smallefjord, c.h. ostenfeld land. the name originated from the wintering party at kulhus during the 1931–34 treårsekspeditionen. canta bræ 71ø-366 (71°58.0´n 25°10.9´w; map 5). glacier on the sw side of sefström gletscher, stauning alper. named by the 1963 cambridge university expedition, the name derives from cante brigge, an ancient name for cambridge. according to myth the university of cambridge was founded by prince cantaber of spain. cantabrigia tinde 71ø-362 (71°55.5´n 25°09.5´w; map 5). moun tain 2780 m high at the head of canta bræ. it was climbed by a cambridge university expedition on 18 august 1963. see also canta bræ. (cantabrigia). cantons-land 72ø (72°40.0´n 27°00.0´w). name used by eugène wegmann during the 1931–34 treårsekspeditionen (wegmann 1935) for the present area of gletscherland, and given for the cantons of switzerland. canutusdal 72ø (72°31.1´n 23°56.3´w). minor valley west of karup elv. the name was used by the 1974 joint biological expedition. canyondalene 74ø-359 (74°32´n 20°05´w). valley system in nw wollaston forland. named during lauge koch’s 1936–38 expeditions by wolf maync and andreas vischer for the canyon-like valleys. (canion dalene.) canyonflod 74ø-360 (74°33.6´n 20°09.9´w). river draining through canyondalene, nw wollaston forland, to enter the sea at albrecht bugt. named by wolf maync and andreas vischer during lauge koch’s 1936–38 expeditions. cap – see also cape, kap and kapp. cap albert de belgique 77ø (c. 77°54´n 19°34´w). this may be a cape in southern hertugen af orléans land, or possibly one of the islands east of hagen ø. it was observed from a great distance by the duke of orléans in 1905, and named after albert 1 [1875– 1934], king of belgium from 1909. cap alf 74ø (74°07.8´n 20°40.3´w). cape on se clavering ø east of dødemandsbugten. the name occurs as c. alf on a sketch map in gustav thostrup’s 1921 logbook (in: møller 1939), and was possibly named after alf trolle [1879–1949], one of the committee of østgrønlandske fangstkompagni. see also kap alf trolle. kapp landmark has also been used. cap blosseville 74ø (74°04.8´n 22°17.0´w). name used for the cape at the se foot of blosseville bjerg at the front of wordie gletscher by karl koldewey’s 1869–70 expedition, the present kap ruth. the name was subsequently transferred to the mountain – see blosseville bjerg. cap de guise 77ø (77°42.5´n 19°11.1´w). alternative name for kap louise on the south side of the mouth of orléans sund. it was named by the duke of orléans in 1905, probably for his cousin jean, duc de guise [1874–1940], who succeeded orléans as pre tender to the french throne. it is used only on one of the folding maps in orléans (1907a). cap deegen 73ø (73°53.1´n 20°56.6´w). name proposed by karl koldewey’s 1869–70 expedition for a cape thought to be on the north coast of hold with hope, but probably corresponding to the present diener bjerg; there is no well defined cape here. named after kammerrichtstrat deegen of leipzig, promoter of the 1873 german west african expedition and a supporter of german arc tic expeditions. (cap degen). cap duc des abbruzzes 78ø (c. 78°20´n 21°20´w; fig. 9). cape or mountain in southern hertugen af orléans land, named by the duke of orléans in 1905 after luigi amedeo abbruzzi. a member of the italian royal family, he was noted for an expedition to franz josef land in 1901 during which a new farthest north record was set on the ice of the arctic ocean. the cape was observed from a great distance, and could not be precisely located by subsequent explorers. cap hélène 77ø (c. 77°19´n 20°02´w; fig. 9). cape on the south side 147 of skærfjorden, sw of kap li, possibly the northern end of the present valdemarsmuren west of slædelandet. named by the duke of orléans in 1905, probably after his grandmother hélène de mecklenbourg-schwerin [d. 1858]. cap holcha 74ø-f34 (74°12.7´n 29°06.8´w). cape on east clavering ø corresponding to the present kap breusing. the name appears as c. holcha on a sketch map in gustav thostrup’s 1921 logbook (in: møller 1939), and was occasionally used by østgrønlandske fangstkompagni. it has also been used in the forms c. holga, kap olga or kap holka (e.g. madsen 1925). cap marie 77ø (c. 77°21´n 19°48´w; fig. 9). cape on the south side of skærfjorden, so named by the duke of orléans in 1905, probably after his wife marie dorothée d’autriche [d. 1932]. the position of the cape could not be definitely fixed by subsequent expeditions, but may have been the present kap li. cap pse. maud 78ø (c. 78°25´n 21°25´w; fig. 9). cape on one of the northern danske øer, named by the duke of orléans in 1905, possibly after princess maud who became queen of norway in 1905. the position of the cape could not be fixed by the 1906–08 danmark-ekspeditionen. cap aase 74ø (74°08.8´n 20°30.1´w). minor cape on se clavering ø west of basaltkap. the name appears as c. aase on a sketch map by gustav thostrup in his 1921 logbook (in: møller 1939). girl’s name. cape – see also cap, kap and kapp. cape beaufoy 74ø (c. 74°30´n 19°20´w). this feature was observed at a great distance by william scoresby jr. in 1822, and may have been a mountain in wollaston forland, possibly huhnerbjerg. it was named after colonel mark beaufoy [1764–1827], a british astronomer and physicist. cape blosseville – see blosseville bjerg. cape bright 74ø (c. 74°37´n 19°00´w). one of the summits of sabine ø, this feature was named by william scoresby jr. in 1822 and placed on his chart north of his kater bay. it was probably named after the physician richard bright [1789–1858], a contemporary of scoresby’s at the university of edinburgh. cape brown mountain 71ø (71°47.1´n 22°26.2´w). name used in a report by säve-söderbergh (1937) for the mountain making up kap brown, the north point of wegener halvø. see also kap brown. cape carnegie 71ø (c. 71°40´n 22°50´w). probably a mountain on wegener halvø, this feature was observed from a great distance by william scoresby jr. in 1822 and could not be identified by subsequent expeditions. it was named in compliment to a much respected edinburgh family. cape crawford 71ø (c. 71°40´n 22°15´w). named by william scoresby jr. in 1822 after an edinburgh friend, the name was intended for a cape on canning land halfway between kap allen and kap fletcher. however, scoresby’s map is difficult to reconcile with modern maps and his cape may have been a mountain west of ålborg fjord. cape hold with hope – see hold with hope. cape krusenstern 71ø (71°36.3´n 22°33.4´w). name given by wil liam scoresby jr. in 1822 to a cape on the west side of the present carlsberg fjord, the present nordenskiöld bjerg. it commemorates the russian navigator adam johann von krusenstern [1770– 1846], who made several notable voyages, including the first rus sian navigation of the world in 1803–06. cape mewburn 72ø (c. 72°12´n 22°09´w). headland on traill ø north of kap moorsom, so called by william scoresby jr. in 1822 after john mewburn, a school friend at whitby who had shared lodgings with scoresby while at the university of edinburgh where he studied medicine. the name seems to have been applied to an insignificant rounding of the coast on the south side of gåsebugt, and has not been used by subsequent explorers. (kap mewburn.) cape read 70ø (70°59.0´n 21°46.0´w). cape on the coast of liver pool land between randers fjord and mariager fjord, the present ravnenæs. the name was proposed by helge g. backlund during the 1931–34 treårsekspeditionen to honour the british geologist herbert harold read [1889–1970]. read was noted for his work in the scottish highlands, especially on granitic rocks, while employ ed by the british geological survey, and as professor at the uni versity of liverpool and imperial college. cape ross 71ø (c. 71°30´n 25°00´w). william scoresby jr. in 1822 reported cape ross as a bold promontory, but was uncertain wheth er it formed part of jameson land or some other distinct region. ryder (1895) said that the cape did not exist at the position indicated, while bay (1896) placed it at a position approximating that of vandreblokken. it is probable that scoresby saw the distant mountains of the southern stauning alper beyond sydkap. the supposed cape was named by scoresby after captain john ross [1777–1856], who had made an important arctic voyage in 1818 to davis strait. (cap ross). cape rossel 73ø (73°08.1´n 23°15.0´w). william scoresby jr. gave this name in 1822 to what he thought was a cape, but was probably a mountain on ymer ø, perhaps the present celsius bjerg. the name, which has not survived, was given out of respect to elisabeth paul edouard de rossel [1765–1829], honorary vice-admiral in the french marine, and first president of the société royale de géographie. cape stufenberg 74ø (74°37.5´n 18°30.3´w). name given to the sw cape of lille pendulum by j.m. wordie’s 1926 expedition, which was named after the mountain above it, karl koldewey’s stufen berg, now terrasseberg. cape syntektite 71ø (71°04.5´n 21°41.4´w). name proposed by helge g. backlund for the present kap buddicom, liverpool land, during the 1931–34 treårsekspeditionen. the name derives from the geology, but was never approved, and occurs on only very few maps (e.g. kranck 1935). capella plateau 73ø (71°04.5´n 21°41.4´w). name given by lauge koch’s 1929–30 expeditions to the plateau area west of margre the dal, corresponding to the present vestreplateau. carbondal 74ø (c. 74°25´n 20°15´w). name used by dunbar (1955) for a valley in western wollaston forland where lauge koch col lected rock samples of carboniferous age. the exact location is uncertain, but it is probably the present sandstensdal, in which flows the river alfred rosenkrantz had called karbon elv. cardiocerasbjerg 74ø-153 (74°28.9´n 20°15.7´w). mountain c. 1680 m high in western wollaston forland, named during the 1931–34 treårsekspeditionen by hans frebold for finds of the fossil ammonite cardioceras. (cardiocerasbjærg.) cardiocerasdal 74ø-92 (74°26.1´n 20°16.4´w). small valley in western wollaston forland draining sw from cardiocerasbjerg into young sund. named by lauge koch’s 1926–27 expeditions as cardioceras valley for the common occurrence of the fossil am monite. cardioceraselv 74ø (74°26.1´n 20°16.4´w). name used by rosen krantz (1932) during lauge koch’s 1929 expedition, for the river flowing in cardiocerasdal, western wollaston forland. cardioceraskløft 70ø-38 (70°44.2´n 25°18.7´w). ravine on the coast of east milne land between charcot havn and kap leslie. the name was used by hermann aldinger during the 1931–34 treårsekspeditionen in the form cardiocerasschlucht or cardio ceras-schlucht, after the fossil ammonite. (cardioceras valley, car dio ceras ravine.) carissima dal 73ø-435 (73°03.3´n 25°13.3´w). valley in east suess land, south of skildvagten, named by silvio eha during lauge koch’s expeditions. as used by eha (1953) the name included the lake and the valleys draining both west (in front of his carissima gletscher) and east into antarctic sund. carissima gletscher 73ø (73°02.9´n 25°16.7´w). name occasionally used by eha (1953) for the glacier sw of niviarsiat which drains 148 southwards into carissima dal. carl heger ø 76ø-20 (76°29.4´n 21°25.0´w; map 4). island in the sw part of dove bugt, named by the 1906–08 danmark-ekspedi tionen as carl hegers ö. probably named by henning bistrup after a member of his family, where the names ‘carl’ and ‘carl heger’ are found (j. løve, personal communication 2009). (karl hegers ø, hegers ø, carl heger island.) carl ritterhytta 76ø (76°07.3´n 19°44.8´w). norwegian hunting station built in 1932 by john giæver’s expedition at kap carl rit ter, on the east coast of ad. s. jensen land. it was originally known as olestua and has also been known as beurmann and ullestuen. (kap carl ritter.) carlsberg dal 71ø (71°25.7´n 22°55.1´w). name used by stauber (1940) for the valley passagen in ne jameson land, which drains into carlsberg fjord. it derives from his work during lauge koch’s 1936–38 expeditions. carlsberg fjord [kangerterajitta itterterterilaq] 71ø-46 (71°25.6´n 22°24.1´w; maps 3, 4). fjord between canning land and liverpool land, first observed by william scoresby jr. in 1822, who thought it connected with hurry inlet. it was mapped by g.c. amdrup’s 1898–1900 expedition, which had the official name ‘carlsbergfondets expedition til øst-grønland’. the carlsberg foundation, which derives its funds from the sales of carlsberg beer and mineral waters, continues to support scientific and cultural activities. (carlsberg fiord, carlsberg-fjord, carlsberg fjorden, karlsbergfjord.) carlsbergfondet land 76ø-111 (76°33.0´n 24°00.0´w; maps 2, 4). part of dronning louise land, south of borgjøkel. named carls berg fondets land by j.p. koch’s 1912–13 expedition, for the most generous single contributor to the expedition’s finances. see also carlsberg fjord. carlshavn 73ø-40 (73°45.8´n 20°27.1´w; map 4). bay in eastern hold with hope, south of home forland. both the bay and the hunting station at the head of the bay built in 1920 were named after the station motorboat carl; the boat was abandoned at bass rock in 1924. norwegian maps from about 1929 used carlshamn for the bay. (carls harbour, karlshavn.) carlshavn 73ø (73°46.3´n 20°28.6´w). danish hunting station at the head of the bay carlshavn, on the east coast of hold with hope. it was built by østgrønlandske fangstkompagni in 1920, manned from 1920 to 1924, and accidently burnt down by norwegian hunters in the autumn of 1927. see also carlshavn. the station has also been referred to as station ‘a’. (karlshavn). carraradal 71ø-373 (71°34.9´n 28°31.0´w). narrow valley in hinks land draining into the head of flyverfjord. named by peter vogt during lauge koch’s 1957 expedition for the outcrops of marble, a tribute to the noted italian marble from carrara. carrick spids 72ø-366 (72°09.3´n 24°47.8´w; map 5). twin rock spires 1970 m high sw of dunottar gletscher in the northern stau ning alper. named by malcolm slesser’s 1958 expedition, which made the first ascent, after carrick castle on loch goil, argyll shire, scotland. (carrick.) caspar spids 71ø-257 (71°55.2´n 23°46.4´w; map 5). mountain about 1450 m high in the se werner bjerge on the east side of søndergletscher. named during the 1953–54 lauge koch expeditions by peter bearth and eduard wenk, and climbed by wenk in 1953. castle 72ø (72°13.5´n 24°39.1´w; map 5). mountain 1830 m high at the head of tårnfjeld gletscher, north stauning alper. first climbed by the 1963 imperial college expedition, and named after the london locality, elephant and castle, originally a smithy which was converted to a tavern in 1760. castor 71ø (71°50.6´n 25°30.8´w; map 5). peak 2520 m high on the sw side of the upper basin of spærregletscher. climbed by karl m. herligkoffer’s expedition on 19 august 1966, it is one of two granite pinnacles, the other which they did not climb being named pollux. the names are derived from the twins of greek mythology, which also gave rise to the names of the stars castor and pollux. castor elv 70ø-181 (70°35.5´n 22°24.3´w). one of a pair of similar rivers in south liverpool land draining west into hurry inlet, the other being known as pollux elv. named during the 1931–34 treårsekspeditionen by laurits bruhn. see also castor. castor glacier 71ø (71°57.5´n 25°41.1´w; map 5). one of two minor tributaries to spærregletscher on its western side, so named by james clarkson’s 1961 expedition. see also castor. german climbing expeditions have used grosse sydney gletscher for the same glacier. castorbjerg 72ø-478 (72°04.2´n 26°16.1´w). mountain north of furesø, nathorst land, opposite polluxbjerg on the south side of the lake. named during lauge koch’s 1954–55 expeditions by hans zweifel together with polluxbjerg for two similar named mountains in the swiss canton of wallis. see also castor. catalinadal 71ø-357 (71°05.0´n 26°50.0´w; map 4). major valley in south renland with several large lakes. named by the 1963 geodætisk institut expedition, at the suggestion of j.v. helk. the valley had apparently been known by this name since the area was photographed during catalina flights by the royal danish air force for the geodætisk institut in 1950. tindernes dal has also been used. cavendish gletscher 72ø-502 (72°05.6´n 25°09.9´w; map 5). gla cier in the stauning alper, draining north to gully gletscher. named by the 1963 cambridge university expedition for the ca ven dish physical laboratory, cambridge, england. cecilia nunatak 72ø-412 (72°30.1´n 27°52.3´w; maps 3, 4). large nunatak west of gletscherland and south of goodenough land. mapped by lauge koch on reconnaissance flights in 1932 during the 1931–34 treårsekspeditionen, and named after the daughter of the british admiral goodenough. see also goodenough land. (cäcilia nunatak, cæcilia nunatak, caecilia nunatak.) celsius bjerg 73ø-27 (73°08.1´n 23°15.0´w; map 4). mountain 1426 m high on eastern ymer ø. named by a.g. nathorst’s 1899 expedition as celsius berg, probably for anders celsius [1701– 1744], a swedish astronomer who was professor in astronomy and mathematics at the university of uppsala from 1729. he was the most noted of three astronomers in the family. this is probably the mountain which william scoresby jr. had called cape russel in 1822 (white 1927). (celsius mountain, mount celsius, celsius fjellet.) centralbjerg 71ø-351 (71°11.9´n 22°53.9´w). mountain 630 m high in east jameson land, west of the head of carlsberg fjord. it was named by john h. callomon during the lauge koch expeditions. centralen 71ø-279 (71°55.0´n 24°03.1´w; map 5). mountain 1370 m high in the werner bjerge between sirius gletscher and aldebaren gletscher, named by peter bearth and eduard wenk during lauge koch’s 1953–54 expeditions. see also centralen. centralen 72ø (72°01.3´n 24°02.3´w). name used by styger (1951) for a mountain between mellem gletscher and østre gletscher, north werner bjerge, the present kolossen. this position for cen tralen is used in a number of climbing reports (e.g. monzino 1966; fantin 1969), but the name is only approved for the position de fined by peter bearth and eduard wenk (see above). centralen – see margarinecentralen. centrumspasset 74ø-348 (74°26.9´n 19°44.6´w). pass at the head of dronning augustadalen in central wollaston forland (centrum = centre). named during lauge koch’s 1936–38 expeditions by wolf maync and andreas vischer (maync 1947). (centrumspas.) centrumsø 80ø-76 (80°10.5´n 22°00.0´w; maps 1, 4; fig. 24). lake in southern kronprins christian land, observed from the air in 1938 by lauge koch. on some maps it has been shown to drain through sødalen and given the name troldsøen (e.g. nielsen 1941; drastrup 1945). it acquired its present name in 1952–53 when it 149 became the natural centre of geological activities after catalina aircraft landed parties here. (centrum sø). cerburus 72ø (72°04.1´n 25°14.1´w; map 5). mountain about 2000 m high between gully gletscher and sefström gletscher, stauning alper. it was climbed by the 1964 zurich expedition, and was so named because it resembled a dog’s head. see also kerberus. charcot bugt – see charcot havn. charcot gletscher 70ø-30 (70°45.0´n 25°46.5´w; map 4). glacier on east milne land at the head of charcot havn. the name appears to have first been used by aldinger (1935) in his report on work during the 1931–34 treårsexpedion, and was named after jean-baptiste charcot [1867–1936]. a french polar explorer and oceanographer, charcot led two expeditions to the antarctic in 1903–05 and 1908–10, and visited the scoresby sund region of east greenland seven times between 1925 and 1936. in 1932 he had transported one of lauge koch’s seaplanes aboard the pourquoi pas? from iceland to scoresby sund. charcot died in the shipwreck of the pourquoi pas? off iceland in 1936. french scientists used glacier chatton for the same glacier. charcot gletscher 73ø (73°02.8´n 29°00.0´w). name used during the 1968 ggu expedition for an e–w-trending glacier dissecting northern charcot land in the inner scoresby sund region (olesen & reeh 1969). use of the name was abandoned when it was found to have been previously given to a glacier on milne land, and this glacier at present has no name. charcot havn 70ø-29 (70°46.8´n 25°23.3´w; maps 3, 4). bay on the east coast of milne land. the name was first used in reports of the work of the 1931–34 treårsekspeditionen in the form charcots harbour (thorson 1934), and commemorates jean-baptiste char cot, whose expeditions had carried out geological work in the vicinity between 1925 and 1936. the name is found on many maps in the form charcot bugt. see also charcot gletscher. chattonbugt has also been used. charcot land 71ø-147 72ø-415 (72°00.0´n 29°00.0´w; maps 3, 5). land area at the head of nordvestfjord between daugaard-jensen gletscher and f. graae gletscher. the name first appears on the 1932 1:1 million scale geodætisk institut map prepared on the basis of aerial observations by lauge koch during the 1931–34 treårsekspeditionen. see also charcot gletscher. (charcots land.) charpentier gletscher 72ø-465 (72°57.9´n 25°56.0´w). glacier in southern goodenough land draining into agassiz dal. named during louise boyd’s 1937 expedition as charpentier glacier after johann von charpentier [1786–1855], a swiss naturalist whose work on glaciers in 1830–40 was closely related to that of louis agassiz. chatham elv 77ø-76 (77°33.8´n 19°12.0´w; map 4). river draining the se part of stormlandet. named during the 1931–34 treårseks peditionen by david malmquist for a friend, gottfrid nordland, usually known as ‘chatham’, who subseqently became headmaster and dean in gällivaare, sweden. chattonbugt 70ø (70°46.8´n 25°23.3´w). bay on east milne land, the present charcot havn. the name was used in the report by parat & drach (1934), who visited the region during j.-b. charcot’s 1933 expedition. it was named after a. chatton, captain of the expedition ship pourquoi pas? in 1932 and 1933. (baie chatton.) chattonkløft 70ø (70°44.5´n 25°29.1´w). gulley sw of charcot havn on east milne land, equivalent to the small valley termed kosmocerasdal by callomon & birkelund (1980). the name was used by parat & drach (1934). see also chattonbugt. chokoladebjerg 73ø-422 (73°22.3´n 25°14.8´w; see also fig. 74). mountain 1010 m high on western ymer ø, north of blomster bugten. the name was given by arthur b. cleaves and ernest f. fox in the course of geological work during john k. howard’s 1933 expedition, originally in the form big chocolate mountain. eha (1953) adopted the name during his geological studies, and it was eventually approved in the present form. the name records the conspicuous deep brown colour of the rocks. chopin dal 76ø-321 (76°42.6´n 23°56.5´w; map 4; fig. 21). valley trending e–w in central dronning louise land between himmer land hede and beethoven dal. one of the names given by the 1952–54 british north greenland expedition after composers, it commemorates frédérik françois chopin [1810–49], a polish musician noted especially for his piano solos and concertos. christian iv gletscher 69ø-33 (69°00.0´n 30°20.0´w; map 3). major glacier draining from geikie plateau southwards to the blosseville kyst. the glacier is said to have been partly mapped by gino watkins, but its extent was first realised during flights by lauge koch in 1933 during the 1931–34 treårsekspeditionen. the name first appeared on maps in the form king christian iv glacier, and commemorates the danish king, christian iv [1577–1648], king of denmark and norway from 1588. he was noted for his establishment of a powerful navy, the foundation of many towns (including christiania, now oslo), and for many fine buildings in copenhagen. christians skær 76ø (c. 76°20´n 19°25´w). skerry east of bælgen, nanok ø, in dove bugt. discovered and so named during the 1932 gefion expedition after one of the danish hunters, christian jensen, who prevented the ship from running into it ( jennov 1935). the name is used in den grønlandske lods (1968). christianshavn 74ø (74°09.9´n 20°11.7´w). danish hunting station built in 1921 at kap mary, eastern clavering ø, by østgrønlandske fangstkompagni, beside a norwegian hunting hut originally built in 1909 (see maryhuset). the station may have been named after christian thielst [1877–1968], who was on the board of østgrøn landske fangstkompagni. the danish station was manned from 1921 to 1923, and was then moved to sandodden, after which the norwegian hut at this location was sometimes referred to by this name. (christians harbour.) christinabjerg 72ø (c. 72°02´n 25°03´w). peak about 2350 m high on the north side of kirkbrae, ne of sefström gletscher, stauning alper. climbed and named by the 1968 scottish expedition. churchill pas 72ø-512 (72°01.8´n 25°01.5´w; map 5). pass between the head of storgletscher and kirkbrae, a side glacier to sefström gletscher. named by the 1963 cambridge university expedition after churchill college, cambridge, founded in 1960 and named after sir winston churchill. see also winston bjerg. (churchill col.) cicero 72ø (72°04.5´n 25°07.4´w). mountain 2400 m high on the east side of cavendish gletscher, northern stauning alper. climbed on 26 july 1984 by sandro pucci’s expedition, and named after the roman orator and statesman marcus tullius cicero [106–43 bc]. cima blonde 72ø (72°08.5´n 25°04.7´w). peak on the ne side of vertebræ, on the north side of gully gletscher, stauning alper. climbed on 29 july 1984 by sandro pucci’s climbing expedition, and probably named for the light colour of the rocks. cima caesar 72ø (72°08.1´n 24°58.9´w). peak wnw of danske tinden, north stauning alper. climbed on 6 august 1984 by sandro pucci’s climbing expedition, and named after the roman general and statesman gaius julius caesar [100–44 bc]. cima di granito 72ø (72°05.0´n 24°39.2´w). name used by guido monzino’s 1963 expedition for glamis borg, a 2200 m granite peak on the sw side of bersærkerbræ. the expedition made the second ascent by a new route. cima est 72ø (72°08.8´n 25°08.9´w; map 5). peak about 2500 m high on the south side of vikingebræ, north stauning alper. first climbed by guido monzino’s 1964 expedition, and probably named after the mountain of the same name in the dolomites, one of the tre cime. cima marco aurelio 72ø (c. 72°07´n 25°07´w). peak on the north side of gully gletscher, north stauning alper. climbed on 2 au gust 1984 by sandro pucci’s climbing expedition, and named after 150 the roman emperor marcus aurelius [ad 121–180]. (m. aurelio). cima ouest 72ø (72°08.9´n 25°10.3´w; map 5). peak about 2400 m high on the south side of vikingebræ, north stauning alper. first climbed by guido monzino’s 1964 expedition, and probably named after the mountain of the same name in the dolomites, one of the tre cime. cima virgilio 72ø (72°04.6´n 25°07.2´w). pinnacle on satans gal leri, the ridge running nne from korsspids, north stauning alper. climbed on 6 august 1984 by sandro pucci’s climbing expedition, and named after the roman poet virgil [70–19 bc]. (virgilio). cirkusbjerg 73ø-77 (73°28.4´n 22°59.5´w). mountain on gauss halvø between karin dal and paralleldal, named by helge g. backlund during lauge koch’s 1929–30 expeditions in the form mt. karboncircus. the original name was found to be geologically misleading and was changed to circus mountain (subsequently cirkusbjerg) at the suggestion of gunnar seidenfaden and helge g. backlund (säve-södergergh 1934). (karboncircus bg.) cirkusdal 74ø-167 (74°20.0´n 20°42.5´w). upper part of djævle kløften, ne clavering ø. so named by arne noe-nygaard and gunnar säve-söderbergh during the 1931–34 treårseks pedi tionen for the bowl-shape of the valley. cirkuselv 73ø-292 (73°57.0´n 22°08.6´w). river in east hudson land draining ne into loch fyne. the name was used by gunnar säve-söderbergh during the 1931–34 treårsekspeditionen, originally as cirkus river, because it drains a cirque or bowl-shaped valley. cirkusgletscher 71ø-256 (71°57.8´n 23°45.6´w). glacier on the east side of the werner bjerge, at the head of blomsterdal. named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk for the arena-like shape. it has occasionally been called escher von der linth gletscher. cirkuskløft 77ø-90 (77°05.3´n 21°34.2´w). ravine east of farve fjældet on the north side of sælsøen. named by eigil knuth’s 1938–39 mørkefjord expedition, probably for its shape. cirque d'acropole 71ø (71°56.0´n 25°58.0´w). name used by the 1968 claude rey expedition for a traverse of the snow domes around their glacier des oubliettes on the west side of prinsesse gletscher, eastern nathorst land. the peaks include l'acropole, dôme de trappeur, dôme de blizzard, dôme des seracs and dôme de leopard. the traverse was named after the acropolis of athens. citadel 71ø (71°40.3´n 25°03.5´w; map 5). mountain about 2000 m high on the north side of mercurius gletscher, south stauning alper. first climbed by james clarkson’s 1961 expedition, and so named because when first seen it towered above all the neighbouring peaks. città di roma 72ø (c. 72°11´n 25°10´w). peak about 2410 m high in the vikingebræ area of the north stauning alper, climbed by g. dionisi’s 1982 expedition. exact location uncertain. it was named after the italian city of rome. city hytta 75ø (75°01.4´n 20°37.9´w). norwegian hunting hut at kap negri on the west side of fligely fjord, built by sigurd tolløf sen’s expedition in august 1932. it was named after gerhard ‘city’ antonsen [1900–1945], a legend among norwegian hunters who spent seven years without a break at moskusheimen. the hut has also been called kap negri hytten. (citystua.) claraiadal 71ø-400 (71°36.5´n 22°59.2´w). valley in the sw part of wegener halvø. named by katherina perch-nielsen during the 1967–72 ggu scoresby sund expeditions for the fossil mussel ‘claraia’, commonly found in the valley. clare fjeld 72ø-503 (72°065.3´n 25°11.9´w; map 5). snow peak 2220 m high on the south side of gully gletscher, stauning alper. climbed on 19 august 1963 by the cambridge university expedition. it was named after clare college, cambridge, founded in 1326 as university hall, and refounded in 1338 by lady elizabeth de clare. (clare.) clare lloyd river 72ø (72°33.7´n 24°00.0´w). name used by the 1974 joint biological expedition for the river in lunedal draining into holm bugt, sw traill ø. it was named after one of the expedition participants. clare’s pingo 71ø (71°59.5´n 23°21.8´w). name used by the 1974 joint biological expedition for a pingo on the south side of kolle dalen, north scoresby land. see also clare lloyd river. clark bjerg 74ø-6 (74°21.8´n 19°14.3´w; map 4). mountain about 400 m high in eastern wollaston forland. it was observed at a distance by william scoresby jr. in 1822, and named cape clark in compliment to john clark, who had married scoresby’s sister mary. the cape was identified as a mountain south of the entrance to dronning augustadalen by the place name committee in about 1935. (clark bjærg.) claudius clavus bjerge 71ø-166 (71°54.5´n 23°12.0´w; map 4). mountain range 900–1100 m high north of ørsted dal, scoresby land. the name was one of a group of names given by the place name committee in 1939, and commemorates the danish cartographer claudius clavus, who prepared some of the earliest maps of greenland. claverhouse 71ø (71°54.6´n 24°52.2´w; map 5). mountain about 2300 m high between storgletscher and gannochy gletscher, central stauning alper. named by the 1968 university of dundee expedition which made the first ascent, probably for john graham of claverhouse [1649–89], 1st viscount of dundee. clavering bukta 74ø (74°14.5´n 20°20.0´w). name used by norwe gian hunters in the 1920s and 1930s for the present kirchenpauer bugt, ne clavering ø (see e.g. white 1927). clavering fjorden 74ø (74°08.0´n 21°53.0´w). name used by norwegian hunters, and on nsiu maps from about 1929, for the sound on the south side of clavering ø now known as godthåb gulf. the 1908–09 floren expedition appears to have been the first to have used the name, although they may have intended it for the present young sund or possibly kirchenpauer bugt, north of clavering ø (see also clavering bukta). (claveringfjorden, clav eringsfjorden, clavering fjord, clavering sund, klaveringsfjorden.) clavering landet – see clavering ø. clavering ø 74ø-78 (74°17.0´n 21°08.0´w; maps 2, 4; fig. 15). large island west of wollaston forland. named by karl koldewey’s 1869–70 expedition as clavering insel after douglas charles clav ering [1794–1827], commander of the griper on the 1823 voyage to this region (sabine 1825; clavering 1830). the west side of the island is separated by a narrow channel from the mainland, which gave rise to reports that it was joined to the mainland and should be called clavering landet (hansen 1912). (clavering island, clave ringöya.) claveringstrædet 74ø-4 (74°31.5´n 19°05.8´w; maps 2, 4). strait between sabine ø and wollaston forland. named by karl koldewey’s 1869–70 expedition as clavering strasse, for douglas charles clavering (see clavering ø). it corresponds approximately to william scoresby’s kater bay. (clavering strait.) cleft island 72ø (72°16.2´n 22°00.7´w; fig. 12). small island off kap young, eastern traill ø, so named by j.m. wordie’s 1926 expedition; officially it has the name ‘rock’. the island has a split appearance. cliff lake 77ø (77°33.3´n 20°49.0´w). lake south of klægbugt, nordmarken. named by the 1987 irish expedition to northern east greenland. cloos klippe 76ø-324 (76°48´n 24°53´w; map 4). cliff on the south side of borg gletscher, central dronning louise land. named by the 1952–54 british north greenland expedition after the german structural geologist hans cloos [1885–1951], professor at the university of breslau 1919–26 and subsequently at the university of bonn. he was a pioneer of granite tectonics. col de furesoe 71ø (71°50.4´n 25°40.2´w; map 5). pass between the heads of prinsessegletscher and borgbjerg gletscher, eastern nathorst land. named by claude rey’s 1968 expedition after 151 nearby furesø. col de scoresby 71ø (71°50.1´n 25°41.9´w; map 5). pass between the heads of prinsessegletscher and borgbjerg gletscher, eastern nathorst land. named by claude rey’s 1968 expedition after william scores by jr. see also scoresby land. col de la tourmente 71ø (71°52.8´n 25°42.0´w). pass between two tributary glaciers in the upper part of prinsessegletscher, eastern nathorst land. named and first climbed by claude rey’s 1968 expedition. col des jaspes 70ø (70°41.5´n 26°02.1´w). col in the mountain range south of charcot gletscher, east milne land, explored by a group from j.b. charcot’s 1933 expedition (parat & drach 1934). prob ably named for the presence of the mineral jasper in the basalts. col des pulkas 72ø (72°00.2´n 24°59.1´w; map 5). high col (2130 m) between kirkbrae and storgletscher, discovered in may 1985 dur ing a w–e crossing of the northern stauning alper. col major – see majorpasset. cold shoulder 72ø (72°04.5´n 24°54.2´w; map 5). peak 2450 m high on the east side of upper gullygletscher, northern stauning alper. climbed and named by the 2007 smc east greenland expedition. colinedal 73ø (73°09´n 25°08´w). valley on sw ymer ø draining southwards, where colin hallenstein located a tungsten-antimony mineralisation while prospecting for nordisk mineselskab (har pøth et al. 1986). colle colosseum 72ø (c. 72°08´n 25°05´w). col 1950 m high between colosseum gletscher and vertebræ, north stauning alper. climbed on 27 july 1984 by sandro pucci’s climbing expedition, and named after the colosseum in rome, one of the most impressive of roman remains. colle genova 70ø (70°03.9´n 23°16.1´w). broad col on torv gletscher, volquaart boon kyst. named by leonardo bonzi’s 1934 expedition, after the italian city. the expedition also used the name ghiacciao genova for the present torvgletscher. colle milano 70ø (70°03.0´n 23°02.0´w). col at the head of milano gletscher, between punta gilberti and punta balestrieri, volquaart boon kyst. named by leonardo bonzi’s 1934 expedition. milan was the point of departure of the expedition. collet bjerg 73ø-316 (73°48.5´n 23°00.0´w). mountain about 1550 m high in central hudson land, ne of ritomsø. named during lauge koch’s 1936–38 expeditions by heinrich bütler after léon wil liam collet [1880–1957], a noted swiss geologist and geomorphologist. (collets bjerg, colletberg, colletbjerg.) coloradodal 71ø-188 (71°33.3´n 23°46.7´w). valley in north jame son land draining nw into ørsted dal. so named by hans stauber during lauge koch’s 1936–38 expeditions because the valley is incised into coloured layered rocks reminiscent of the grand canyon, colorado. coloradodal hytten 71ø (c. 71°34´n 23°58´w). hut in northern jameson land built in july 1983 for grønlands miljøundersøgelse where the rivers draining coloradodal and major paars dal meet at qilerneq. it was used by a group studying and marking musk ox. colosseum gletscher 72ø (c. 72°07´n 25°04´w). glacier on the north side of gully gletscher, stauning alper. named by the 1984 sandro pucci climbing expedition after the colosseum in rome. see also colle colos seum. coltart 71ø (71°58.0´n 25°01.7´w; map 5). summit 2395 m high in the upper reaches of sefström gletscher. climbed by the 1998 scottish mountaineering club expedition, and so named for a shape like a lobster claw. combe d’argent 71ø (71°54.8´n 25°54.8´w). tributary glacier on the west side of prinsessegletscher, eastern nathorst land. named by claude rey’s 1968 expedition, perhaps for the colour (combe d'argent = silver comb). commandment peak 71ø (71°07.3´n 26°14.9´w). high point, 2127 m high, on the ice cap south of edward bailey gletscher and east of catalinadal, renland. climbed and named by the 2007 west lancashire mountaineering group expedition. concordia 71ø-334 (71°43.3´n 25°05.4´w; map 5). confluence of several glaciers in central bjørnbo gletscher, stauning alper, forming a broad level area. so named by john hunt’s 1960 expedition after similar glacier confluences in the swiss alps. concordia fjeld 73ø-416 (73°58.3´n 28°04.4´w). nunatak in arnold escher land north of a confluence of glaciers. named during lauge koch’s 1951 expedition by hans r. katz. concordia plads 72ø-444 (72°38.6´n 27°49.6´w; maps 3, 4). con fluence of glaciers north of cecilia nunatak. so named by eugène wegmann during the 1931–34 treårsekspeditionen after the similarly named glacier confluences in the swiss alps. (konkor dia platz). cône des eboulis 71ø (71°59.8´n 25°58.6´w). mountain about 1600 m high west of the front of prinsessegletscher. named and first climbed by claude rey’s 1968 expedition. consolation point 71ø (71°09.2´n 26°18.7´w). summit 1914 m high south of edward bailey gletscher, renland. climbed and named by the 2007 west lancashire mountaineering group expedition. constable pynt 70ø-133 (70°44.5´n 22°35.8´w; maps 3, 4). low peninsula on the west side of hurry inlet, the northernmost point in the fjord reached by william scoresby sr. in 1822. it is also the location of the airport known as constable pynt [nerlerit inaat], built in 1985. the name originated from william scoresby jr. as point constable, and appears to be placed on scoresby’s map due west of the fame øer (fig. 3). although the latter position was retained on per dusén’s map (nathorst 1901), on the map in am drup (1902a) it is placed 7 km sw of the fame øer, the present site and that used on nearly all maps since amdrup. named after archibald constable [1774–1827], bookseller and publisher, who had published several of scoresby’s books. the airport was constructed at constable pynt to serve the oil and gas exploration centred on jameson land, and subsequently largely replaced mestersvig airfield. (konstabel pynt.) continental banke 76ø (c. 76°45´n 15°00´w). offshore bank east of germania land. the name appears to have first been used by the 1906–08 danmark-ekspeditionen (e.g. johansen 1912). conus i, conus ii – see kegle i, kegle ii. copeland fjord 74ø-123 (74°15.0´n 22°02.0´w; map 4). n–strending fjord on the west side of clavering ø, named by lauge koch’s 1929–30 expeditions. ralph copeland, astronomer and physicist of karl koldewey’s 1869–70 expedition, made observations in the vicinity in october 1869, and had discovered the connection between rudi bugt and copeland fjord. see also copeland gletscher. (copelands fjord.) copeland gletscher 74ø-321 (74°36.9´n 22°11.0´w). glacier on the sw side of tyrolerdal, payer land, named by louise boyd’s 1937 expedition after ralph copeland, who accompanied julius payer to the vicinity of this glacier in 1869–70. pasterze was used for this glacier on the 1932 geodætisk institut 1:1 million scale map, and on some maps the names of copeland gletscher and kløft gletscher are interchanged. see also copeland fjord. (cope land glacier.) copeland gulf 74ø (74°08.0´n 21°53.0´w). name used by rodahl (1946) for the present godthåb golf, south of clavering ø, an ex tension of copeland fjord. see also copeland fjord. copelandshytten – see kap copeland hytten. cordulaspids 71ø (71°58.7´n 24°54.5´w; map 5). mountain 2430 m high on the west side of upper storgletscher, central stauning alper. climbed and named after a living person by the 2007 smc east greenland expedition. corrugated roof ridge 73ø (75°23.9´n 27°18.9´w). name used in a report by the 1972 university of dundee expedition for the ridge on the north side of haredalen, ne frænkel land. it was climbed on 20 august, and has a series of regular ravines grooving its side. cotton peak 73ø (73°32.7´n 26°01.1´w). peak 1979 m high on the 152 south side of grejsdalen, andrée land. climbed by the 2007 army boreal zenith expedition. courier passet 71ø (71°54.1´n 24°56.9´w; map 5). easy pass be tween dalmore glacier and gannochy gletscher, central stau ning alper. explored and named by the 1968 university of dundee expedition. courtauld bjerg 74ø-144 (74°17.6´n 22°28.6´w). mountain 1255 m high west of clavering ø. the name was originally used by lauge koch’s 1929–30 expeditions in the form courtauld land for the area of which the present courtauld bjerg is the highest point. the name commemorates augustine courtauld [1904–59], a noted brit ish arctic explorer. he took part in wordie’s 1926 and 1929 cambridge expeditions to east greenland, but is best known for his five months’ isolation at a meteorological station on the inland ice during the 1930–31 british arctic air route expedition. (courtaulds bjerg ). craig øer 72ø-7 (72°23.5´n 22°20.7´w; map 4). islands in mount norris fjord. they were named the craig islands by william scores by jr. in 1822 after a much respected episcopalian clergyman of edinburgh. scoresby evidently intended the name to apply to eight islands, including those close to the north side of mount norris fjord, but the name is usually used in a more restricted sense for the four islands in the centre of the fjord. (craig öer, craigöya). crescent pas 72ø-510 (72°03.7´n 24°55.8´w; map 5). col or pass between the heads of gully gletscher and storgletscher, stauning alper, first reached by the 1961 bangor expedition. it may have been given its name by the 1963 cambridge university expedition. (crescent col). crescent tind 72ø (72°03.6´n 24°57.2´w; map 5). summit about 2450 m high on the west side of crescent pas, at the head of gully gletscher, stauning alper. climbed and so named by the 1996 norwegian stauning alper expedition. crinoid bjerg 70ø-401a (70°30.0´n 23°04.7´w). minor hill in south jameson land on the west side of muslingeelv. it was originally named in the form crinoid mt. by hermann aldinger (1935) during the 1931–34 treårsekspeditionen for finds of fossil crin oids, although the name was first approved in 1972 at the suggestion of the 1967–72 ggu scoresby sund expeditions. crossopterygian ravine 73ø (73°30.6´n 23°25.4´w). ravine on the south side of the west end of sederholm bjerg, gauss halvø. the name was used by johansson (1935), and records his finds of vertebrate fossils during the 1931–34 treårsekspeditionen. (crossop terygie-ravinen.) culross 71ø (71°40.3´n 25°12.5´w; map 5). mountain about 2067 m high on the south side of jupiter gletscher, southern stauning alper. named by james clarkson’s 1961 expedition after culross palace, fife, scotland. curie klippe 76ø-310 (76°57.6´n 25°11.2´w; map 4). cliff south of admiralty gletscher in dronning louise land. one of the names given by the 1952–54 british north greenland expedition for notable scientists, it commemorates pierre and marie curie who with henri becquerel were jointly awarded the nobel prize in 1903 for their research into radioactivity. czoks topp 71ø (71°53.7´n 25°05.2´w; map 5). mountain about 2490 m high on the north side of roslin gletscher, between the two branches of the minor glacier valhallbreen. it was climbed by the 1996 norwegian stauning alper expedition, and so named after andrzej czok [1948–1994] a polish climber and colleague who died climbing in the himalayas. d d. eglin spire 72ø (72°04.3´n 24°45.4´w; map 5). peak about 2500 m high at the head of bersærkerbræ, between bersærker tinde and royal peak, stauning alper. it was climbed and so named by the 1985 i.m. marsh college expedition. d’aunay bugt 69ø-15 (69°00.0´n 25°32.0´w). deep bay or fjord south of kap ryder on the blosseville kyst. the name was adopted by g.c. amdrup’s 1898–1900 expedition from a map made by jules de blosseville in 1833, where it appears as baie d’aunay. raoul d’aunay was second in command of the la lilloise, blosseville’s ship during his 1833 voyage to east greenland. (d’aunay-bugt, d’aunay fjord.) daedalus 71ø (71°40.7´n 25°07.1´w; map 5). mountain 2040 m high between jupiter gletscher and mercurius gletscher, southern stauning alper. named and first climbed by james clarkson’s 1961 expedition. daedalus was the mythical greek architect said to have built the labyrinth for king minos of crete. dagfinvika 72ø (72°41.3´n 22°03.5´w; fig. 14). bay on south geo graphical society ø, west of kap mcclintock. used only on nsiu maps (lacmann 1937), the name was given for the norwegian painter and sculptor dagfin werenskiold [1892–1977] who participated in the 1931 nsiu expedition to the area. see also weren skioldflya. dagmar havn 76ø-78 (76°40.7´n 18°52.7´w). small bay on the ne coast of store koldewey, named by the 1906–08 danmark-ekspe di tionen as dagmars havn. possibly named after the 640-ton cor vette dagmar on which j.p. koch, a member of the expedition, trained for entry into the navy in 1887. (dagmar harbour). dagmar havn hytten 76ø (76°38.9´n 18°46.9´w). norwegian hunting hut built in september 1938 by the norsk–franske polarekspedisjon near dagmar havn, ne store koldewey. it has also been called øresundshytten, but this is misleading as it is located on the strait named lille bælt and not that named øresund. dagny-bankerne 75ø (c. 75°35´n 17°45´w). offshore banks ne of shannon. according to jennov (1935) they had been known by this name since the dagny noted depths of 40–50 fathoms here in 1919, and was sunk over the banks in 1920. the dagny was a schooner which carried the first party of danish hunters to east greenland for the østgrønlandske kompagni in 1919. alf trolle considered danmarksbankerne a more appropriate name, because the 1906–08 danmark-ekspeditionen had found the north side of the banks in 1906. the 1968 edition of den grønlanske lods uses the form dagny banke. daguerrefjellet 74ø (74°21.9´n 21°06.9´w). snow summit about 1585 m high on north clavering ø between ortlerspids and høj nålen. so named on nsiu maps of lacmann (1937) after louis jacques mandé daguerre [1789–1851], the frenchman who in vented the daguerreotype. dahl skær 74ø-106 (74°09.3´n 20°18.4´w). small island off eastern clavering ø. the name first appears on a sketch map in the 1921 logbook of gustav thostrup (møller 1939) in the form dahlsskær. so named after kai r. dahl, who as a journalist for the danish newspaper berlingske tidende sailed in this area with the teddy in 1921 and 1923. (dahls skerry, dahl island, dahl insel.) dahl skær hytten 74ø (74°09.6´n 20°18.5´w). norwegian hunting hut built in the summer of 1948 for hermann andresen’s expedition immediately north of dahl skær, eastern clavering ø. the hut was enlarged by sirius in 1970. (dahls skær hytten). dahlisfjell – see dæhlis fjell. dalføret 73ø-434 (73°00.4´n 25°47.5´w). broad pass at the crest of nanortalikdal in suess land (dal = valley, føret = the passage). dalheim 73ø (73°30.4´n 23°40.1´w). norwegian hunting hut on the south side of gauss halvø at the mouth of paralleldal. built in august 1930 by arktisk næringsdrift, who used the name troms dal for the valley because of a similarity with the region in norway. dalhytten 74ø (74°30.2´n 20°37.9´w). danish hunting hut at the east end of store sødal, ne of zackenberg. built by nanok in august 1938, it has also been known as blæsenborghytten. dalmore glacier 71ø (71°54.0´n 25°00.0´w; map 5). glacier on the north side of roslin gletscher. explored and named by the 1968 university of dundee expedition. 153 dalmore junior 71ø (71°52.5´n 25°05.3´w; map 5). mountain about 2140 m high on the north side of roslin gletscher, west of dalmore glacier. the name was used by the cambridge university expedition which climbed it on 27 july 1970. dalskuta 74ø (74°13.1´n 21°04.4´w). mountain 1454 m high on south clavering ø. the mountain lies at the north end of skræl lingedalen (skut = a prominent cliff or rock-wall). the name is used only on nsiu maps (lacmann 1937). dalstrøget 73ø (73°32.2´n 24°50.6´w). valley on the west side of geologfjord, andrée land, possibly identical with tillitkløft. the name is used in den grønlandske lods (1968). daltærskel 80ø-121 (80°12.0´n 21°30.0´w; map 4; fig. 24). local ity at the east end of centrumsø, from which sæfaxi elv drains eastwards (tærskel = threshold). named during operation groundhog 1960. damelv 70ø-165 (70°45.6´n 22°25.4´w). river in south liverpool land draining west into hurry inlet, so named during the 1931–34 treårsekspeditionen by laurits bruhn because it drains a small lake (= dam). damesten 73ø-655 (73°32.5´n 24°28.0´w). boulder on the southernmost flat peninsula of strindberg land. named during the 1931–34 treårsekspeditionen by th. johansen after the glaciertransported boulder known by this name on fyn, denmark. dammen 72ø-90 (72°03.3´n 25°28.5´w; map 5; fig. 38). embay ment at the south end of alpefjord almost completely dammed by sefström gletscher and gully gletscher. named during the 1931– 34 treårsekspeditionen by ove simonsen. in the past the glaciers formed a complete dam and a series of terraces record water levels up to 60 m above present sea level (sugden 1962). damslottet 72ø-267 (72°03.9´n 25°40.9´w; maps 4, 5). mountain west of dammen, at the corner between innermost alpefjord and the east end of furesø. named by john haller following explorations during lauge koch’s 1954 expedition (slottet = the castle). this mountain is probably identical with mitternachspitze, climbed in 1971. daneborg 74ø-278 (74°18.2´n 20°13.2´w; maps 2, 4). this name was originally given to the nordøstgrønlands slædepatrulje base at sandodden, built in 1944 with assistance of usa forces (daneborg = the danes castle), and is the name now used for the present sirius headquarters. at the end of the war the base was taken over as an icao weather station, with significant extensions in 1947, 1952 and 1961. sirius daneborg, headquarters of slædepatruljen reso lute, was established a few hundred metres south of the weather station in 1951. the weather station closed in the summer of 1975, and the buildings were taken over by the sledge patrol. daneborg has been noted for many years for its large colony of eider ducks (70 pairs in 1964) which nest between the tethered dog teams, and are thus protected from foxes (rosenberg et al. 1970). see also sand odden and karina. danevirke 72ø-210 (72°11.9´n 23°45.9´w; map 5). ridge se of noret, north scoresby land. named by prospecting teams associated with lauge koch’s 1948–49 expeditions, after the complex of earthworks in sydslesvig between trene and slien, the oldest dating from ad c. 750. daniel bruun land 76ø-112 77ø-118a (76°53.0´n 21°52.0´w; maps 2, 4). land area between sælsøen and borgfjorden, named by j.p. koch’s 1912–13 expedition as daniel bruuns land. daniel bruun [1856–1931], a captain in the danish navy and author of several books on the arctic, had assisted ludvig mylius-erichsen in planning the 1906–08 danmark-ekspeditionen. daniel schmidtfjellet 74ø (74°23.3´n 21°09.3´w). mountain 1400 m high on north clavering ø. named after daniel schmidt [b. 1902], who undertook photogrammetric work on the detailed nsiu maps of clavering ø and geographical society ø (lacmann 1937). danmark havn 76ø-35 (76°45.7´n 18°41.3´w; map 4). sheltered harbour in south germania land, so named by the 1906–08 dan mark-ekspeditionen after the expedition ship danmark, which winter ed here. the danmark, a 377-ton barque built in sunder land in 1855, had previously sailed under the names sir colin campbell, steamer, and under norwegian ownership as the magdalene. in 1906 it was purchased for the expedition and renamed danmark. the name has commonly been used in the form danmarks havn (with genetive ‘s’), and it is this form that is most often seen on maps. skibshavn has also been used. the icao weather station danmarkshavn (spelt as one word) is at the north side of the harbour. (danmark harbour, danmarkshamnen, dan mer kurhöfn.) danmark ø [ujuaakajiip nunaa] 70ø-67 (70°30.0´n 26°15.0´w; maps 3, 4). island in the inner part of scoresby sund, named dan marks ø by carl ryder’s 1891–92 expedition for the kingdom of denmark (fig. 7). the expedition wintered in hekla havn on danmark ø. ragnvald knudsen occasionally used heklalandet for the island in his diaries (knudsen 1890). (dan marks ö, danmark island, ile de danemark, dänemark-i.) danmarks monumentet 76ø-49 (76°56.1´n 20°59.3´w). im pres sive mountain buttress between inner mørkefjord and pustervig, fig. 38. looking eastwards towards the stauning alper, with dammen and krabbegletscher in the foreground. the john haller photograph collection, geus archive. stauning alper dammen k rab begletsch er 154 daniel bruun land, so named during the 1906–08 danmarkekspeditionen. according to charles poulsen (1991), and tho strup (2007) it received its name from the pattern of light and dark rocks in the steep east face, which resembled the monogram of the danish king, christian ix. (monumentet, monumentfjæld, monu ment fjeldet, monumentum daniæ, danmarks monument, danmer k ur-minnisvarðis, chr. d. 9. monumentet.) danmarks pynt 70ø (70°26.8´n 28°14.6´w). name used by gulløv (1991) for the east peninsula of hekla havn where the over-wintering station was erected (j. løve, personal communication 2010). danmarksbankerne – see dagny-bankerne. danmarkshavn 76ø-198 (76°46.2´n 18°40.2´w; maps 2, 4). weath er station at danmark havn, the site of the original base of the 1906–08 danmark-ekspeditionen. the modern icao weather and radio station around østerelv was established in 1948 and is still in use. fischer (1983) reported it as manned by 11 men and comprising 15–16 houses. relief ships normally sail to danmarks havn with supplies in alternate years, but ice not uncommonly prevents their arrival. danmarkshavnhuset 76ø (76°46.2´n 18°41.1´w). name often used for the base of the 1906–08 danmark-ekspeditionen at danmark havn, which was originally known as villaen. it was briefly used by the 1909–12 alabama expedition, and repaired and taken over in 1919 as a hunting station by østgrønlandske fangstkompagni. from 1929 it was used by nanok. the graves of two hunters, robert frørup and hans nielsen, who died of scurvy after the dagny sank in 1920 and provisions failed to arrive, are behind the house, as is the memorial to the three lost members of the 1906–08 danmarkekspeditionen. the house is now known as danmarksminde (fisch er 1983; p.s. mikkelsen 1994, 2008). (danmarkshavn sta tion, danmarks minde.) danmarksminde – see danmarkshavnhuset. dannebrogsfjeldene 76ø-145 (76°40.0´n 25°18.4´w). range of mountains and nunataks in sw dronning louise land, east of revaltoppe. named by j.p. koch’s 1912–13 expedition as danne brogs fjeldene after the danish flag, the dannebrog, said to have fallen from the sky in 1219 during a battle between danes and esto nians at reval, an old nordic name for the capital of estonia. (dannebrogs tinder, dannebrogs fjall.) danske roseneath – see mønstedhus. danske villa 73ø (73°53.2´n 21°52.5´w). danish hunting hut on the west side of loch fyne, immediately south of strømmen, built by nanok in september 1950. it has also been known as strømmen hytten. a norwegian hut known as norsk villa is on the opposite side of the fjord. danske øer 77ø-147 78ø-48 (78°07.0´n 19°00.0´w; maps 1, 4). island group in jøkelbugten. named by john haller during lauge koch’s 1956–58 expeditions, to commemorate the work of the 1906–08 danmark-ekspeditionen. dansketinden 72ø-266 (72°07.5´n 24°57.3´w; maps 4, 5; fig. 27). highest mountain in the stauning alper, 2842 m high, situated between the heads of vikingebræ, gullygletscher and bersærker bræ. it was first climbed by john haller, wolfgang diehl and fritz schwarzenbach on 5 august 1954. the second ascent was made by guido monzino’s 1964 expedition. (dansketinden = the danes peak.) darien 71ø (71°50.2´n 25°24.7´w; map 5). small snow peak about 2400 m high on the divide between the heads of bjørnbo gletscher and spærregletscher, stauning alper, first climbed by james clarkson’s 1961 expedition. darien was the name formerly applied to the entire district of panama, where, from a high peak, vasco nuñez de balboa first saw the pacific ocean in 1513. some view this peak as identical with muhldorfer spids. darien pass 71ø (71°50.3´n 25°24.5´w; map 5). pass between the upper part of bjørnbo gletscher, and the head of spærregletscher, stauning alper. the name is used by bennet (1972), and derives from nearby darien. daudmannsvågen – see dødemandsbugten. daudmannsøyra 74ø (74°07.3´n 20°54.5´w). coastal stretch of døde mandsbugten on south clavering ø, where numerous inuit houses and graves occur (daudmann = dead man). the name was reported in 1930 as used by norwegian hunters, and subsequently incorporated in nsiu charts (lacmann 1937). danish hunters used the roughly equivalent term dødemandstomten. (daumanns øyra, daudmannsöyra.) daugaard-jensen gletscher 71ø-65 (71°50.0´n 28°47.0´w; maps 3, 4). large glacier between hinks land and charcot land, and one of the most productive in east greenland. the name first appears on the 1932 1:1 million scale geodætisk institut map prepared on the basis of aerial observations by lauge koch during the 1931–34 treårsekspeditionen. it was named after jens dau gaard-jensen [1871–1938], a danish administrator who was in spector of nw greenland from 1900, and director of grønlands styrelse from 1912–1938. he was president of the 1931–1934 expedition committee, and koch notes that he constantly followed the work of the expedition with interest. (daugård-jensen gletscher.) david gray hytten – see kap david grayhytten. davy sund 71ø-140 72ø-12 (72°04.5´n 22°40.0´w; maps 3, 4). wide inlet connecting to the nw with kong oscar fjord. william scoresby jr. named davy’s sound in 1822 for sir humphry davy [1778–1829]. davy was a noted chemist, president of the royal society from 1820–1827, and most remembered for his invention of the miners safety lamp. nathorst (1901) suggested latitude 72°10´n as the limit of davy sund, the approximate present limit, while white (1927) suggested the limit ought to be carried as far as kap peterséns and the haslum øer. (davys sund, davis sund, david sund, davysund). davy sund hytten 71ø (71°57.0´n 22°44.1´w). norwegian hut on the south side of davy sund, nw of kap biot, built in august 1930. it has also been known as biot-stua and villa. (david sund hytten). de dødes bjerg – see dødemandstoppene. dead lake 77ø (77°34.4´n 20°54.9´w). large lake sw of klægbugt, nordmarken. named by the 1987 irish expedition to northern east greenland. deceit bugt – see lumskebugten. deichmann fjord [pukkitsivakajiip oqqummut kangertiva] 69ø-22 (69°49.0´n 23°14.0´w; maps 3, 4). fjord sw of manby halvø on the northern blosseville kyst. named by g.c. amdrup’s 1898–1900 expedition after henrik deichmann [1871–1939], entomologist, physician and ornithologist on the expedition. he had also taken part in carl h. ryder’s 1891–92 expedition as zoologist, and subsequently practised as a doctor in west greenland. delta øen 72ø (72°15.5´n 23°57.2´w; map 5). name occasionally used on maps in the 1950s for the large area between two branches of tunnelelv, that before mestersvig airfield was built appeared as an island during the melt when rivers were high. deltadal 72ø-297 (72°05.3´n 23°58.0´w; map 5). wide, flatbottomed valley draining into mesters vig, marked by a meandering network of streams and deposits of sand and mud. the name was adopted by the place name committee from a suggestion by peter bearth and eduard wenk during lauge koch’s 1953–54 expeditions. sieburgerdal, siborgdal and storedal have been used for the same valley. demos glacier 71ø (71°41.3´n 25°03.5´w; map 5). minor glacier on the sw side of bjørnbo gletscher, stauning alper. the name ap pears on some of the maps of james clarkson’s 1961 expedition and in bennet (1972). possibly named after deimos, a small moon of mars. den ny hytte 76ø (76°30.3´n 20°14.2´w). hut built by danmarks havn weather station personnel in 1966 on an island in dove bugt, ne of licht ø. the name means ‘the new hut’. den store nanuták 76ø (76°40.0´n 24°20.0´w). this was said by 155 trolle (1909) to be the original name for dronning louise land, the extensive region of large and small nunataks west of l. bistrup bræ and storstrømmen. the name translates as ‘the large nunatak’. den lille sø 76ø (76°46´n 18°42´w). name used by trolle (1909) for a small lake near danmark havn, the present drikkevandsø. the name translates as ‘the small lake’. dendritgletscher 69ø-31 (69°35.0´n 25°38.0´w; map 3). large, many branched glacier system in northern christian ix land, draining east to blosseville kyst. the name was given for the striking dendritic pattern first observed by lauge koch on flights in 1933 during the 1931–34 treårsekspeditionen. dentdal 72ø-518 (72°30.0´n 23°50.8´w). small valley on western traill ø draining north into karupelv. named by geoffrey halli day following botanical work during the 1961 leicester university expedition and 1971 northern universities expedition. denti della norsketinde 72ø (72°08.6´n 25°03.2´w). peak about 2500 m high in the northern stauning alper, north of norske tinden, climbed and so named by g. dionisi’s 1982 expedition. depotelv 71ø-45 (71°12.2´n 23°11.1´w). river in jameson land rising on the slopes of fossilbjerget and treford bjerg, and flowing west into hall bredning. named by g.c. amdrup’s 1898–1900 expedition as depot elven during the first exploration of jameson land in august 1900 by otto nordenskjöld and henrik deich mann. depotfjeld 80ø-32 (80°10.5´n 16°55.3´w; map 4). mountain in se holm land. so named by eigil nielsen during the 1938–39 mørke fjord expedition because depots were made at its foot. the 1906–08 danmark-ekspeditionen had sometimes referred to this mountain as mallemukfjeld, but the latter name was established by eigil knuth’s 1938–39 mørkefjord expedition as referring to a nearby mountain. depotgletscher 80ø-33 (80°10.9´n 16°46.1´w). glacier between depotfjeld and mallemukfjeld in se holm land. named by eigil nielsen during the 1938–39 mørkefjord expedition. see also depotfjeld. depotkulle 77ø-87 (77°04.3´n 20°26.4´w). mountain about 420 m high on the north side of sælsøen. named during the 1938–39 mørkefjord expedition by paul gelting who had made a depot at the foot of the mountain in june 1939 when exploring sælsøen. depotnæsset 77ø-104 (77°33.7´n 18°54.0´w; maps 1, 2, 4). east ern most cape of stormlandet. so named by eigil knuth’s 1938–39 mørkefjord expedition because both the 1906–08 danmarkekspeditionen and the mørkefjord expedition had made depots here. svend sølver found a depot here in april 1939 laid out by gaston micard’s 1938–39 expedition. depot ryg – see smalleryg. depotskæret 76ø-12 (76°02.1´n 19°48.6´w). small island off se ad. s. jensen land. so named by the 1906–08 danmark-ekspedi tionen, probably for the prominent inuit cairn built as a marker for a large meat depot (thostrup 1911). den grønlandske lods (1968) uses depotskærene for a group of nine islands between trums ø and kap beurmann. (depotskjæret, depot reef.) depotsten 73ø (c. 73°56´n 21°53´w). rock on the east side of loch fyne, north of strømmen. the rock had been used as a depot marker, and the name occurs as a botanical reference locality in reports of the 1931–34 treårsekspeditionen (gelting 1934). depotø [immikkertaa] 71ø-51 (71°38.6´n 22°30.0´w). island on the south side of nathorst fjord, so named by lauge koch’s 1926–27 expeditions as depot island because supplies were cached here during his sledge journeys. derry 71ø (71°41.3´n 24°36.5´w). mountain 1480 m high north of the front of bjørnbo gletscher, the present snekuppel. climbed during john hunt’s 1960 expedition, and named after derry lodge, aberdeenshire, where young men selected for the expedition were trained. det lille rød hus – see washburns hus. devaux plateau 70ø-34 (70°44.4´n 25°27.7´n). plateau south of charcot havn, east milne land, so named during the 1931–34 treårsekspeditionen by hermann aldinger after j. devaux, a member of the 1933 charcot expedition which visited the area. he was drowned in the wreck of the pourqoui pas? in 1936. (de vaux-plateau.) devoldhalvøya 72ø (72°54.1´n 22°00.0´w). broad peninsula on geographical society ø, north of cambridge bugt. so named on the nsiu maps of lacmann (1937), after the brothers finn, halvard and joakim devold. see also devoldhytta, hallvardvatnet, finnvatnet and joakimpasset. devoldhytta 73ø (73°17.7´n 24°26.0´w). norwegian hunting hut on the north side of dusén fjord, ymer ø, west of zoologdalen. it was built in 1929 for arktisk næringsdrift by olav kjelbotn and hal vard devold, and named after halvard ophus devold [1898– 1957]. a norwegian telegraphist, he worked at meteorological stations in finnmark, svalbard and jan mayen between 1920 and 1926, and helped to found arktisk næringsdrift for whom he worked as a hunter from 1929–1932. in 1931, either on his own initiative or at the suggestion of activists in norway, he took part in the annexation of eirik raudes land, an action which led to the dispute between norway and denmark over the sovereignty of east greenland. he was appointed secretary of nsiu in 1940, but was captured by the us coast guard while leading a relief expedition to the norwegian hunting stations in east greenland. he spent the war years in a camp on the isle of man. (devold, devold hytte.) devon canyon 73ø (73°40.4´n 24°35.3´w). name used by poulsen (1937) for a narrow ravine 1 km south of gunvor bjerg, strindberg land. it was given for the rocks of devonian age. devon hills 73ø (73°53.9´n 22°11.2´w). name used by lauge koch in 1930 for the 900 m high mountains between the nørlund alper and nordhoek bjerg, ne hudson land, corresponding to the present passagehøje. they were originally named for the presumed occurrence of devonian rocks. helge g. backlund suggested the name be discontinued in favour of his passage hills (now pas sagehøje) when the rocks proved to be carboniferous in age. (devonhaugane, devon hill, devon høje.) devondal 71ø-402 (71°35.9´n 22°41.7´w). valley on south weg ener halvø, draining into nathorst fjord. so named by katharina perch-nielsen during the 1967–72 ggu scoresby sund expeditions because the southernmost outcrops of devonian rocks in east greenland are found here. devonpas 74ø-399 (74°14.0´n 24°11.7´w). pass in ole rømer land leading eastwards to vibeke gletscher. named by john haller during lauge koch’s 1956–58 expedition, after occurrences of devonian rocks. diadem 71ø (71°58.9´n 24°57.9´w; map 5). name used by hans gsellman’s 1957 expedition for a three-peaked mountain about 2400 m high east of granta bræ, stauning alper, because of the manner in which the peaks caught the sun. according to bennet (1972) the west peak was subsequently climbed in 1963 by a cambridge university party and is now known as downing fjeld. the second ascent of two of the peaks was made by a 1968 party led by donald bennet. fantin (1969) and bennet (1972) give different positions for this peak, the uncertainty arising from the quality of gsellman’s original maps. diamond peak 71ø (71°49.6´n 25°01.5´w; map 5). peak about 2150 m high on the south side of roslin gletscher, south stauning alper. climbed by the 1982 sheffield university expedition. diannsketinden 72ø (72°07.1´n 24°58.7’w). peak about 2532 m high on the spiky ridge south of dansketinden, stauning alper. so named by the 1996 scottish mountaineering club expedition, although after their return they discovered it had been climbed two months earlier by the 1996 norwegian stauning alper expedition and named tårnet. the norwegian party estimated a height of 156 2310 m. dickens bjerg 76ø-335 (76°23.5´n 26°20.3´w; map 4). prominent mountain in sw dronning louise land. one of the names given by the british north greenland expedition 1952–54 for novelists, it commemorates charles dickens [1812–1870], generally regarded as the greatest english novelist. it was climbed by the lancaster university expedition in may 2000. dickson fjord 72ø-402 (72°50.0´n 27°00.0´w; map 4; see also fig. 52). fjord between suess land and gletscherland. named by a.g. nathorst’s 1899 expedition for robert dickson [1843–1923] who had made contributions to the expedition’s finances. (dicksons fjord.) didrik pining bjerge 71ø-173 (71°40.7´n 23°32.7´w; map 4). mountain range up to 966 m high west of fleming fjord. the name was one of a group of names given by the place name committee in 1939 to replace proposals by hans stauber. it was given for didrik pining, an admiral and administrator in iceland whom christian i sent to greenland in 1476. diener bjerg 73ø-52 (73°53.1´n 20°56.6´w). mountain 800 m high in northern hold with hope, named by lauge koch’s 1929–30 expeditions in the form mt. diener. it commemorates carl diener [1862–1928], an austrian stratigrapher and vertebrate palaeontologist noted for his studies of triassic strata. rocks of triassic age make up the mountain. (dienerfjellet, dieners bjerg.) diener river 73ø (73°54.9´n 20°57.1´w). name used by eigil niel sen during the 1931–34 treårsekspeditionen for a river draining diener bjerg, northern hold with hope (nielsen 1935). dijmphna sund 79ø-29 80ø-2 (80°07.0´n 18°00.0´w; maps 1, 4; fig. 24). sound north of hovgaard ø and south of holm land and lynn ø. named by the 1906–08 danmark-ekspeditionen for the steamer dijmphna, which captained by andreas peter hovgaard became trapped in the kara sea in 1882–83 during an attempt to assist two dutch expedition ships, and lost its screw. see also kap maria dijmphna. (dymphna sund.) dinosaur 71ø (71°40.1´n 25°17.1´w; map 5). highest summit, about 1900 m high, of an impressive row of rock peaks sw of ursus minor gletscher, south stauning alper. first climbed by james clarkson’s 1961 expedition, and named for the profile of the ridge that resembled a dinosaur’s back. dinosaurus elv 70ø (70°31.2´n 22°37.9´w). river on the west side of hurry inlet, near the mouth of which an alleged dinosaur footprint was found by nikolaj hartz in july 1900 during g.c. amdrup’s 1898–1900 expedition. the name was used by rosenkrantz (1934) who expressed doubt as to the identification of the find. (dino saurus river.) dinosaurus klöft 70ø (70°30.7´n 22°37.1´w). name used by rosen krantz (1934) for the ravine west of hurry inlet in which dino saurus elv flows, the present quppaalakajik. dipperne 81ø (81°20.8´n 14°06.5´w). nunataks at the nw margin of kilen, kronprins christian land. the name was given for a sighting of a bird of the dipper family, and is found on a coloured geological map of kilen printed in 1991. disa gletscher 73ø-712 (73°10.3´n 28°22.3´w). minor glacier on the north side of nordenskiöld gletscher, named by j.m. wordie’s 1929 expedition as disa glacier and traversed on the approach to petermann bjerg. the name was approved in 1951 at the suggestion of john haller. (disas gletscher, disagletscher.) diskordansdal 71ø-394 (71°37.4´n 22°54.6´w). valley on wegener halvø. named by katharina perch-nielsen during the 1967–72 ggu scoresby sund expeditions for a well-exposed discordance between two geological formations. dislokationsdal 74ø (74°21.0´n 20°33.1´w). valley on east clavering ø marked by a tectonic dislocation or fault. the name was used by maync (1949). diver loch 72ø (72°15.6´n 23°58.6´w). name used by the 1974 joint biological expedition for a very small lake in the hills west of nyhavn, near mestersvig airfield, and named after the red-throated diver. in some descriptions, two lakes are distinguished, diver north loch and diver south loch. djævleborg 73ø-677 (73°32.9´n 26°18.5´w; map 4). mountain about 1800 m high in central andrée land, on the south side of grejsdal. so named during lauge koch’s 1949–51 expeditions by john haller, because the mountain can be likened to a castle (= borg), and is sited at the entrance to djævlekløft. djævlehjørnet 73ø-685 (73°29.9´n 26°41.1´w). mountain about 2000 m high on the north side of the west entrance to djævle kløften in sw andrée land. named during lauge koch’s 1949–51 expeditions by john haller. djævlehånden 70ø-438 (70°35.0´n 29°34.0´w). flat ice plateau on the north side of paul stern land, formed by the confluence of five glaciers. it is an exposed and windy place where the wind follows the five glacier fingers (djævlehånden = the devil’s hand). named by laurent jemelin during the 1967–72 ggu scoresby sund expeditions. djævlekløften 73ø-642 (73°33.2´n 26°23.2´w; map 4). narrow valley in andrée land, connecting central rendalen with the head of grejsdal. so named by ove simonsen during the 1931–34 treårsekspeditionen because of its wild and threatening appearance. djævlekløften 74ø-113 (74°20.2´n 20°34.5´w). deep valley on east clavering ø, named by lauge koch’s 1929–30 expeditions, apparently for its forbidding appearance. (djævleklöft, djaevlakløften, djevleklöften.) djævlekløfthytten 74ø (74°20.0´n 20°27.8´w). danish hunting hut on the north side of the mouth of djævlekløften, ne clavering ø, built by nanok in august 1930. djævlespalterne 73ø-421 (73°25.0´n 30°30.0´w). area of extensively crevassed glaciers along the margin of the inland ice, west of frænkel land. hans r. katz encountered large numbers of wide crevasses here during his journey with motor-sledges in 1951 (djævlespalterne = the devil’s crevasses). djævleøen 76ø-164 (76°23.3´n 20°24.5´w; map 4). island in west dove bugt. so named by the 1932 gefion expedition because of its association with teufelcap (= devil’s cape) and hestefoden (= horse’s hoof ) on the same island. dobbeltglacier valley – see gletscherdal. dobbeltskæret 76ø (76°47.3´n 18°23.5´w). skerry off the east coast of germania land, south of syttenkilometernæsset. the name is found in thostrup’s (2007) account of the 1906–08 danmark-eks pe di tionen (j. løve, personal communication 2009). dobbelttop 74ø-81 (74°58.0´n 20°08.4´w). mountain 1090 m high on ne kuhn ø, named by karl koldewey’s 1869–70 expedition as doppelgipfel for its two summits. dobbelttoppen 71ø-31 (71°02.7´n 21°56.1´w). mountain 1040 m high south of storefjord in liverpool land. named by william scoresby jr. in 1822 as double mount for its two summits. it is similar to but slightly lower than kirken on the north side of storefjord. (doppelberg.) dobbeltvigen 71ø-94 (71°41.1´n 22°17.6´w). enclosed bay on the west coast of canning land. named during the 1931–34 treårs eks peditionen by arne noe-nygaard in the form doubletvigen, for its two indentations. (doubletvig.) dobbeltøerne 79ø (79°22.0´n 18°43.8´w). two small islands on the south side of nioghalvfjerdsfjorden, the present eli knudsen øer. the name was used by eigil nielsen, who passed by the islands on a sledge journey in june 1939. tvillingøer has also been used. doctordalen – see dronning augustadalen. doggerelv 71ø-192 (71°17.6´n 24°00.0´w; map 4). river in jame son land draining south into fegin elv. named by hans stauber during lauge koch’s 1936–38 expeditions for the age of the rocks (dogger stage of the jurassic period). dolerite point ponds 71ø (71°51.3´n 22°54.2´w). name used in an 157 ornithology report of the 1963 british east greenland expedition (hall & waddingham 1966) for several small lakes in lower ørsted dal, scoresby land. they were named for outcrops of dolerite. dolezalfjellet 72ø (72°55.7´n 23°00.0´w). mountain on central geo graphical society ø, corresponding to the present tørvestakken. used on the nsiu maps of lacmann (1937), the name was given for eduard doležal [1862–1955], an austrian, and one of the leading developers of photogrammetric techniques. dolken 70ø-452 (70°27.3´n 29°20.9´w). mountain 1810 m high in paul stern land. so named by w.e. adrian phillips during the 1967–72 ggu scoresby sund expeditions for its knife-sharp ridge (dolk = knife). dollar 71ø (71°40.0´n 25°11.1´w; map 5). mountain 2085 m high on the south side of jupiter gletscher, south of culross, south stauning alper. first climbed by james clarkson’s 1961 expedition, and perhaps named after the small scottish town near castle campbell, best known for its academy. dolomitdal 74ø-165 (74°22.9´n 20°35.8´w). valley on ne clav ering ø. the name was used by arne noe-nygaard and gunnar säve-söderbergh during the 1931–34 treårsekspedi tionen, be cause of the occurrence of dolomite. (dolomittal.) dolomitpynt 73ø-566 (73°31.4´n 24°41.7´w; map 4). cape in east andrée land nw of kap weber. named by christian poulsen during lauge koch’s 1929 expedition as dolomite point, for the occurrence of dolomite. dombjerg 74ø-64 (74°33.0´n 20°48.0´w; map 4). snow-capped mountain about 1200 m high south of lindeman fjord and north of store sødal. named domberg by karl koldewey’s 1869–70 expedition, possibly for the alpine mountain of similar name. (mt domberg.) dom brava – see dombravahytten. dombrava, dombravap – see dumbrava, dumbravap. dombravadal 70ø (70°37.5´n 22°17.3´w). name briefly in use in the 1930s for gubbedal, liverpool land, which contains the localities dumbravap imia and dumbrava. dombravahytten 70ø (70°36.8´n 22°25.9´w). name used until about the 1950s for the hut which constantin dumbrava built on the east side of hurry inlet at the locality known as dumbrava. the inhabitants of scoresbysund today use a two word version of the name, dom brava, for the hut on the same site. (dumbravahytten.) dome 71ø (71°55.0´n 24°55.5´w; map 5). mountain on the ridge between storgletscher and dalmore glacier, central stauning alper. named by the 1968 university of dundee expedition, which made the first ascent in august of that year. dôme charcot 70ø (70°32.0´n 21°44.2´w). ice cap about 680 m high in south liverpool land, equivalent to the present hvidefjeld [apuseeq]. the french international polar year expedition 1932– 33 had determined the thickness of the ice at 50–70 m, and the name is used on maps in several of their reports (e.g. rothé 1941). it was named after jean-baptist charcot [1867–1936], most noted for his polar explorations. he led french expeditions to the antarctic in the française in 1903–05 and the pourquoi pas? in 1908–10, and later a series of expeditions to east greenland in the pourquoi pas? see also charcot land. dôme de l'envoi 71ø (71°50.6´n 25°46.1´w). snow dome about 2400 m high on the west side of prinsessegletscher, eastern nathorst land. named and first climbed by claude rey’s 1968 expedition. dôme della norsketinde 72ø (c. 72°08´n 25°05´w). peak about 2500 m high in the north stauning alper, near norsketinden, climbed by g. dionisi’s 1982 expedition. dôme des séracs 71ø (71°55.5´n 26°00.2´w). snow dome about 2650 m high on the west side of prinsessegletscher, eastern nathorst land. first climbed by claude rey’s 1968 expedition and named for the crevasses. dôme du blizzard 71ø (71°56.5´n 26°00.5´w). snow dome about 2500 m high on the west side of prinsessegletscher, eastern nat horst land. named and first climbed by claude rey’s 1968 expedition. dôme du leopard 71ø (71°55.0´n 25°57.7´w). snow dome about 2600 m high on the west side of prinsessegletscher, se of dôme des séracs, eastern nathorst land. named and first climbed by claude rey’s 1968 expedition. dôme du trappeur 71ø (71°56.5´n 21°56.7´w). snow dome about 2500 m high on the west side of prinsessegletscher, eastern nat horst land. named by claude rey’s 1968 expedition. domkirken 72ø-205 (72°11.0´n 24°01.1´w; map 5). mountain 1025 m high on the nw side of store blydal, north scoresby land. named by prospecting teams associated with lauge koch’s 1948–49 expeditions for its shape (domkirken = cathedral), most impressive as seen from the site of minebyen. domkirken 73ø (73°32.7´n 20°29.7´w). danish hunting hut at the mouth of glommen about 4 km north of kap broer ruys, hold with hope, built by nanok in september 1945. the name derives from a 3–4 m high tower built onto the hut, which due to heavy snowfall is the only means of access in winter. it has also been known as broer ruys nord. (kirkehytten.) donau passet 71ø (71°50.6´n 25°21.3´w; map 5). pass on the south side of the head of roslin gletscher, leading to the head of bjørnbo gletscher. so named by karl m. herligkoffer’s 1966 expedition after the river donau (danube), the major river which rises in the schwarzwald of southern germany. (donnau pass.) dortes kulmine 70ø (70°27.7´n 22°14.5´w). coal seam about 50 cm thick at the mouth of brudelv, on the west coast of rosenvinge bugt, south liverpool land. it was found by an eight-year-old greenlandic girl (dorte) in 1925, and has been worked periodically. it is now exhausted. it is also known as aamarsuit and ikkaa lissat. doseths fjell 71ø (71°54.9´n 25°06.1´w; map 5). mountain about 2590 m high on the north side of roslin gletscher, stauning alper. it was climbed by the 1996 norwegian stauning alper expedition, and so named after hans christian doseth [1957–1984], a prominent norwegian climber who died climbing in the himalayas. double ravine – see western upper terrace. doumer høj 70ø-370 (70°29.1´n 21°57.3´w). point on the east side of scoresbysund where a memorial to paul doumer was erected by the french international polar year station 1932–33. p. doumer [1857–1932], a mathematician, journalist and politician, was president of france when assassinated in 1932. he was j.b. charcot’s closest friend, and a large portrait of doumer had a place of honour above the mess table in the pourquoi pas? see also ker doumer. the monument was described in 1933 as comprising six whitepainted fuel drums piled on top of each other (nyholm-poulsen 1985). dove bugt 76ø-6 (76°36.0´n 20°00.0´w; maps 2, 4). extensive bay west of store koldewey, bounded to the north by germania land and to the south by ad. s. jensen land. so named dove bai by karl koldewey’s 1869–70 expedition, after the german physicist and meteorologist heinrich wilhelm dove [1803–79]. a prominent scientist he was professor at the university of berlin, and from 1849 director of the prussian meteorological institute (j. løve, personal communication 2010). koldewey’s usage was restricted to the extreme nw part of the present bay. the bay has, somewhat speculatively, been identified with the breidifjòrdr of the icelandic sagas (tornøe 1944). (dove bay.) downing fjeld 71ø-358 (71°58.8´n 25°00.1´w; map 5). snow mountain about 2500 m high south of granta bræ, stauning alper. climbed by the 1963 university of cambridge expedition, and named after downing college, cambridge, founded in 1800 with the proceeds of the estates of sir george downing. (downingfjeld.) draba sibirica elv 71ø-378 (71°06.3´n 23°26.7´w; map 4). river in jameson land draining west to hall bredning. so named by geoffrey halliday following botanical work during the 1961 158 university of leicester expedition, after a species of whitlow-grass. drach kløft 70ø-31 (70°44.8´n 25°34.1´w). ravine se of charcot havn, east milne land, named during the 1931–34 treårsekspedi tionen by hermann aldinger as drach-schlucht. pierre drach, a scientist at the university of paris, was a member of j.b. charcot’s 1933 expedition that visited this region. dragneset 72ø (72°45.8´n 21°58.6´w). peninsula in eastern geo graphical society ø. the name was used on the nsiu maps of lacmann (1937), and derives from the norwegian word ‘sjødrag’ (= swell of the sea). dragøyane 72ø (72°48.7´n 21°57.1´w). small islands in cambridge bugt, off eastern geographical society ø. so named on the nsiu maps of lacmann (1937), the name derives from the norwegian (see dragneset.) draugen 73ø (73°47.1´n 20°16.7´w). skerry off kap kraus in home forland, northern hold with hope. used on an nsiu map (1932a), the name is a norwegian dialect word for a ghost, often a headless evil spirit which appears as a warning of death. dreiecks plateau 74ø (74°45.3´n 20°43.4´w). triangular plateau about 508 m high in th. thomsen land, on the west side of fligely fjord. the name was used by vischer (1943) in a report on work during lauge koch’s 1936–38 expeditions (dreieck = triangle). dreieselbjerg 72ø-308 (72°00.4´n 23°53.7´w; map 5). mountain 1442 m high in the north werner bjerge, named by peter bearth and eduard wenk during lauge koch’s 1953–54 expeditions. it was climbed by a party of three in 1953, a not particularly nice climb of an unimpressive peak. the climbers were ‘the three donkeys’ (= drei esel). dreikant 72ø (72°01.2´n 25°04.2´w; map 5). mountain about 2400 m high on the ne side of sefström gletscher, stauning alper. named for the shape of the mountain, a three-sided pyramid, and first climbed by hans gsellman’s 1957 expedition. dreikonigsgrat 71ø (71°47.8´n 25°26.4´w; map 5). ridge with three peaks between the heads of borgbjerg gletscher and orion glet scher. probably first climbed and named by the 1977 schwäbische stauning alper expedition. dreispitz 72ø (72°11.8´n 25°11.9´w; map 5). three rock peaks, about 2000 m high, on the north side of vikingebræ, north stau ning alper. they were traversed by hermann huber’s 1968 expedition. (trespids.) dresdner spids 71ø (71°55.2´n 25°23.5´w). mountain 2580 m high on the north side of duart gletscher, stauning alper, the present duart borg. climbed by karl m. herligkoffer’s expedition on 17 august 1966, and named after the town of dresden, germany. mont saussure has also been used. (dresdner bjerg.) dreverspids 71ø (71°52.6´n 24°57.4´w; map 5). mountain 2210 m high between dalmore glacier and gannochy glacier, central stauning alper. first climbed by the 1968 university of dundee expedition and named after harold irving drever [1912–75], professor at the university of st. andrews. drever had visited west and north-west greenland nine times and had developed a passionate interest in greenland inuit culture. (drever.) driftwood valley 73ø (73°09.0´n 25°50.0´w). small valley on the ne coast of suess land, east of scheele bjerg. the name is used only in the archaeological report of mci. johnson (1933) describing his work during j.m. wordie’s 1929 expedition. drikkevandsø 76ø-251 (76°46.5´n 18°42.6´w). small lake near danmark havn, se germania land. the name was used during the 1906–08 danmark-ekspeditionen (lundager 1912), as the lake was the source of the expedition’s drinking water. trolle (1909) referred to the same lake as den lille sø. it may be identical with skibssø. drillinge 71ø (71°53.1´n 25°34.4´w; map 5). mountain 2560 m high between hecate glacier and the upper part of spærregletscher, stauning alper, with three conspicuous granite pinnacles (drillinge = three barrelled rifle). it was climbed by karl m. herligkoffer’s expedition on 23 august 1966. it has also been called grosse kederbacher spids. dritte weisse – see tredie hvide. dritten lagergipfels 72ø (72°03.8´n 25°15.5´w). temporary name (dritten lager = camp 3) used by hans gsellman’s 1957 expedition for a 2500 m high mountain on the north side of sefström glet scher, stauning alper. they later called it sonnblick spids (koglbauer 1965). dromledome 81ø (81°15.7´n 13°54.7´w). hill in nw kilen, kron prins christian land, with a dome-shaped geological structure formed in sandstone. soil creep leads to movement of large sandstone slabs that can be heard at long distances. the name is found on a coloured geological map of kilen printed in 1991 (pedersen 1991). dronning augustadalen 74ø-58 (74°25.5´n 19°22.4´w). prom inent valley in eastern wollaston forland. named by karl kolde wey’s 1869–70 expedition as königin augusta thal, after maria luise augusta katharina [1811–1890], queen of prussia and german empress. she had made a substantial donation to the expedition finances. norwegian hunters have occasionally referred to the valley as doktordalen, due to a misinterpretation of the abbreviation ‘dr.’ for ‘dronning’. (königin-augusta-thal, drottning augustas dal, dr. augusta dal, queen augusta valley, augusta dalen.) dronning louise land 75ø-86 76ø-110 77ø-135a (76°40.0´n 24°20.0´w; maps 2, 4). extensive region west of dove bugt com prising several very large and numerous small nunataks. named by the 1906–08 danmark-ekspeditionen as dronning louises land for louise [1851–1926], the swedish wife of the danish king frederik viii. trolle (1909) occasionally referred to it as den store nanuták, which he states was its original name. (queen louise land.) dronning margrethe ii land 75ø-113 (75°40.0´n 21°00.0´w; maps 2, 4; see also fig. 59). land area with a southern boundary at ardencaple fjord and bredefjord and a northern boundary at bessel fjord, equivalent to the present nørlund land together with wollaston forland. it was named after queen margrethe ii of denmark on the occasion of her 50th birthday, 16 april 1990. dronningestolen 76ø-130 (76°31.0´n 25°00.0´w; map 4). moun tain in sw dronning louise land between kursbræ and pony gletscher, named by j.p. koch’s 1912–13 expedition. the mountain was climbed by members of the expedition on 29 april 1913 and provided a magnificent view of all of dronning louise land. named for the association (dronningestolen = queen’s throne), and the well-known locality with the same name at møns klint, denmark. (dronningestol, drotningarstóll.) drumglas 71ø (71°58.7´n 24°52.8´w; map 5). summit 2330 m high on the west side of upper storgletscher, central stauning alper. climbed and named by the 2007 smc east greenland expedition. drumglas beag 71ø (71°59.4´n 24°53.0´w; map 5). peak 2060 m high on the west side of upper storgletscher, central stauning alper. climbed and named by the 2007 smc east greenland expedition. dryasdal 77ø-124 (77°11´n 23°30´w; map 4). valley in north dronning louise land. named by the british north greenland expedition 1952–54 for the occurrence of the plant ‘dryas octo petala’ in what is described as a pleasant, green and flower-filled valley. drygalskifjellet 74ø (74°25.2´n 21°04.0´w). mountain 1500 m high on north clavering ø. so named on the nsiu maps of lac mann (1937) after erich von drygalski [1865–1949], a german geophysicist and geographer who was professor in berlin from 1899 to 1906. he took part in expeditions to west greenland in 1891–93, to the antarctic in the gauss 1901–03, and to spits bergen with count zeppelin in 1910. see also kap drygalski. drømmebjerg 80ø-74 (80°16.1´n 21°27.0´w; map 4; fig. 24). 159 mountain ne of centrumsø, kronprins christian land. named during lauge koch’s 1952–53 expeditions by erdhart fränkl, who had planned to visit the area because of its promising geological relationships, but was forced to return to base; his plans remained but a dream (= drømme). drømmebugten 72ø-60 (72°11.2´n 22°35.5´w; map 4). deep bay on se traill ø. it was first seen by the ‘jägmästeren’ on a.g. nathorst’s 1899 expedition who thought it to be a significant new fjord. discovery of its somewhat restricted extent gave rise to many jokes, and its name as drömbukten (= dream bay). (drombugten, dream bay, drømbugt, drombukta.) drømmetinde 72ø-294 (72°56´n 29°05´w). nunatak summit about 2500 m high on the west side of nordenskiöld gletscher. named during lauge koch’s 1953 expedition by john haller. after a long and difficult traverse across several large glaciers, the party was prevented by bad weather from completing their exploration of this nunatak region. ascent of this peak, the highest mountain, thus remained a dream. dråbegletscher 80ø-48 (80°35.3´n 19°29.0´w). hanging glacier on the east side of innermost ingolf fjord. so named draabegletscher by eigil nielsen during the 1938–39 mørkefjord expedition be cause it resembled a round, clear glass drop (= dråbe), which had hardened halfway down the mountain (nielsen 1941). duart borg 71ø-311 (71°55.2´n 25°23.5´w; map 5). mountain 2583 m high on the ne side of duart gletscher, south stauning alper. named by malcolm slesser’s 1958 expedition after duart castle in the isle of mull, scotland, a 13th century stronghold of the lords of the isles, and now the home of the chiefs of clan maclean. it was first climbed by the 1964 zürich expedition that named it mont saussure. it has also been called dresdner spids. duart gletscher 71ø-310 (71°54.8´n 25°27.5´w; map 4). glacier in the south stauning alper, a branch of spærregletscher sw of duart borg. first traversed by malcolm slesser’s 1958 expedition, and named duart glacier. duart–roslin col 71ø (71°53.2´n 25°21.0´w; map 5). broad flat col at the heads of duart gletscher and roslin gletscher, stauning alper. duck lake 76ø (76°25.2´n 18°45.0´w). lake on store koldewey where samples were taken for radiocarbon age determinations and phytoplankton studies (cremer et al. 2005, 2008). duck pond 72ø (c. 72°14´n 22°54´w). name used by the 1974 joint biological expedition for a small lake at the se end of mestersvig airfield. dudhope 71ø (71°54.6´n 24°54.0´w; map 5). mountain between storgletscher and gannochy gletscher, central stauning alper. named by the 1968 university of dundee expedition, which made the first ascent. dukkegletscher 73ø-649 (73°51.5´n 25°40.0´w). small glacier in ne andrée land, draining into geologfjord. so named by th. johansen during the 1931–34 treårsekspeditionen, probably because of its small size (dukke = doll). dumbrava 70ø-179 (70°36.8´n 22°25.9´w). locality on the east coast of hurry inlet where constantin dumbrava, a rumanian scientist, built a house without permission in 1930, with the intention of trading with the greenlanders. he was picked up by the godthaab and taken back to europe in 1931, and the house was taken over by scoresbysund municipality and used for hunting. the name was recorded by the 1955 geodætisk institut name registration. contemporary accounts of the incident record his name as ‘dombrava’, a spelling carried over into the place names and still found on many maps, although it was officially corrected in 1967. (dombrava.) dumbravahytten – see dombravahytten. dumbravap imia 70ø-177 (70°37.0´n 22°23.3´w). river draining gubbadal, entering hurry inlet at dumbrava. recorded during the 1955 geodætisk institut name registration, the name translates as ‘dumbrava’s water’. (dombravap imia.) dumbravap kangileqitaa 70ø-170 (70°38.9´n 22°27.8´w). minor cape on the east coast of hurry inlet north of dumbrava. recorded during the 1955 geodætisk institut name registration, the name translates as ‘dumbrava’s inner cape’. (dumbravap kanileqitâ, dom bravap kanileqitâ.) dumbravap kanileqitâ – see dumbravap kangileqitaa. dump pool 72ø (c. 72°13´n 23°54´w). name used by the 1974 joint biological expedition for a small lake se of mestersvig airfield near the dump. dunderdalen 72ø (72°02.8´n 23°09.3´w). valley on the nw side of antarctic havn, the present eneboerdal (dunder = thunder, rumble). the name is found on norsk søkort 511, published in 1937. dunholm [immikkeerterajivit] 69ø-24 (69°55.0´n 22°40.0´w). small island ne of steward ø on the north blosseville kyst. named by g.c. amdrup’s 1898–1900 expedition for the down of nesting eider ducks (dun = down). numerous eiders were noted here by n. hartz on 30 july 1900. (dunholme.) dunken 74ø-286 (74°16.5´n 21°49.7´w). mountain on west clav ering ø whose top resembles a square petrol can (= dunk). the name originated from the wintering party at eskimonæs during the 1931–34 treårsekspeditionen. dunlin swamp 72ø (71°59.8´n 23°10.9´w). name used by the 1974 joint biological expedition for the coastal marsh on the sw side of inner antarctic havn, ne scoresby land, where many dunlins were observed. dunlin valley 72ø (72°15.5´n 23°57.5´w). name used by the 1974 joint biological expedition for a minor valley in the hills west of nyhavn, near mestersvig airfield. named for the dunlin. dunne fjæld 70ø (70°37.8´n 22°43.2´w). minor summit 700 m high on the west side of hurry inlet between muskusoksekløft and astartekløft. so named by hermann aldinger during the 1931–34 treårsekspeditionen (aldinger 1935). dunottar bjerg 72ø-365 (72°09.5´n 24°51.1´w; map 5). mountain 2524 m high on the west side of bersærkerbræ, north stauning alper. first climbed by malcolm slesser’s 1958 expedition, and named after dunottar castle, kincardinshire, scotland, a spectacular fortress dating largely from the 14th century, now a ruin. the second ascent was made in 1967 by toni gobbi’s expedition. (dunottar.) dunottar gletscher 72ø-364 (72°08.6´n 24°43.5´w; map 5). gla cier on the west side of bersærkerbræ, north of dunottar bjerg, north stauning alper. named dunottar glacier by malcolm sles ser’s 1958 expedition. dunskjold col 72ø (72°10.8´n 24°50.8´w; map 5). name used by the 1982 sheffield university expedition for the col between dunottar gletscher and skjoldungebræ, stauning alper, situated between achnacarry and elisabethsminde. it was first climbed from the dunottar gletscher side. dunvegan toppene 72ø-367 (72°07.4´n 24°31.7´w; map 5). moun tain summits 1894 m high between bersærkerbræ and skel bræ, north stauning alper. first climbed by malcolm slesser’s 1958 expedition, and named after dunvegan castle, isle of skye, scotland, seat of the chiefs of clan macleod since at least 1200. in an early report of the expedition it went under the name of garbh bheinn. (dunvegan.) durham klippe 76ø-334 (76°21´n 24°52´w; map 4). cliff on the north side of budolfi isstrøm, south dronning louise land. the name was given by the 1952–54 british north greenland expedition and commemorates the university of durham, founded in 1832, from which two of the expedition members, hal lister and peter taylor, had graduated. durin 73ø (73°29.1´n 22°14.3´w). one of the peaks of troels-lund bjerg in the giesecke bjerge. so named on an nsiu map (1932a) for a dwarf in old nordic mythology. see also dvalin. 160 dusén bjerg 70ø-126 (70°58.1´n 22°37.4´w). mountain in east jameson land, ne of the head of hurry inlet. named by alfred rosenkrantz during lauge koch’s 1926–27 expeditions as dusén mt., after per karl hjalmar dusén [1855–1926], the botanist, cartographer and photographer on a.g. nathorst’s 1899 expedition. dusén was the first to map around the head of hurry inlet. see also dusén fjord. (duséns bjerg.) dusén fjord 73ø-30 (73°14.3´n 24°00.0´w; maps 3, 4). e–w fjord almost dividing ymer ø. named by a.g. nathorst’s 1899 expedition after per dusén, who was the first to observe the mouth of the fjord. his mapping work in the two and a half weeks of nathorst’s expedition is regarded as his life’s greatest scientific achievement. see also dusén bjerg. lauge koch’s greenlandic assistants in 1927, karl and tobias, reported the fjord to be considerably longer than first thought. it was first fully explored by lauge koch and by nsiu in 1929. (dusen fjord, duséns fjord, dussinfjorden). dusens fjordhytten 73ø (73°10.6´n 23°08.3´w). hut south of the mouth of dusén fjord, west of kap wijkander, built by arktisk næringsdrift in august 1929. it was known originally under the name kikut, and later occasionally as steffensens hytte. dvalin 73ø (73°28.8´n 22°09.9´w). one of the peaks of troels-lund bjerg in the giesecke bjerge. so named on an nsiu map (1932a) after a dwarf in old nordic mythology. see also durin. dværgarvedal 71ø-379 (71°54.9´n 23°29.8´w). valley draining from the east flanks of the werner bjerge northwards into blom sterdal. named by geoffrey halliday following botanical work during the 1961 university of leicester expedition, after a plant of the carnation family. dværgfjorden 80ø-56 (80°46.3´n 14°15.7´w; map 4). small fjord on the east coast of amdrup land, south of sophus müller næs. so named by eigil knuth during his 1938–39 mørkefjord expedition because of its small size (dværg = dwarf ). dybedal 74ø-191 (74°10.4´n 20°55.9´w). deeply incised tributary valley to skrællingedalen on south clavering ø (dybe = deep). the name was used originally as a botanical reference locality in reports of the 1931–34 treårsekspeditionen (gelting 1934). (dybental.) dybendal 73ø-58 (73°47.9´n 22°43.2´w; map 4). valley in hudson land draining east into stordal, so named during the 1931–34 treårsekspeditionen by th. johansen because of the high, steep valley sides. material for an intended hut was transported into the valley by nanok in march 1952, but with the cessation of nanok’s activities was never built. dyndvulkan – this designation (= mud volcano) is occasionally en countered on published maps (e.g. kempter 1961) for the periglacial ice-cored sand mounds known as pingos (see also pingo dal). dyndvulkanen – see vulkanhytte. dyraelv 73ø-150 (73°29.2´n 21°02.3´w). river in south hold with hope, first named on the 1932 nsiu map (nsiu 1932a; fig. 13) in the form dyra. probably named for the abundant cast antlers of reindeer (dyra = dyr = animal). dyrdalen 73ø (73°29.2´n 21°02.3´w). valley in south hold with hope in which dyraelv flows. the name appears in this form on the 1932 nsiu map (nsiu 1932a; fig. 13), and also in occasional ornithology reports (e.g. bird & bird 1941). dyrfaret 73ø (73°19.0´n 24°48.9´w). name used for a norwegian hunting hut on the north side of innermost dusén fjord (nsiu 1932c), built by arktisk næringsdrift in september 1930. it has also been known as trangen and strømhytta. (dyrstien, dyrstein.) dyrfjellet 73ø (73°33.8´n 21°15.3´w). ridge in the southern tågefjeldene, hold with hope, equivalent to part of the present ravnebjerg. the name appears on the 1932 nsiu map (nsiu 1932a; fig. 13), and may have been named for the abundant cast antlers of reindeer. dyrhö ( full name = dyrhögda) 73ø (73°32.9´n 21°13.7´w). moun tain 1041 m high at the se end of dyrfjellet, hold with hope. the name is found only on the 1932 nsiu map (nsiu 1932a; fig. 13). dyrskolten 73ø (73°39.8´n 21°19.0´w). mountain in the central tågefjeldene, hold with hope. the name appears only on an nsiu map (1932a). dæhlis fjell 71ø (71°54.8´n 25°06.5´w; map 5). summit about 2570 m high on the north side of roslin gletscher. it was climbed by the 1996 norwegian stauning alper expedition, and so named after finn dæhli [1955–1984], a prominent norwegian climber who died climbing in the himalayas. døde bræ 70ø-436 (70°28.0´n 29°12.0´w; map 4). glacier north of paul stern land draining into vestfjord. so named during the 1967–72 ggu scoresby sund expeditions to scoresby sund by ole olesen because studies showed it to be stationary (død = dead). døde slette 73ø-385 (73°37.0´n 25°22.0´w; map 4). ice plateau in eastern andrée land, north of grejsdalen. named by erdhart fränkl during lauge koch’s 1948–50 expeditions for its extremely desolate nature, ‘where not even flowers grow’. (dødesletten gletscher). dødehundebræ 74ø-240 (74°04.9´n 25°39.0´w). glacier in north strindberg land draining into granitsø. it was named as de døde hundes bre by arne høygaard and martin mehren in 1931, because they shot five of their dogs here after crossing the inland ice from west to east. dødemandsbugten 74ø-248 (74°06.8´n 20°53.6´w). bay on se clavering ø. the name appears in the form daudmannsvågen on an nsiu map (1932a), and like the term daudmannsøyra for the coastal stretch, may have been in use earlier by norwegian hunters. at this site there are 43 inuit winter houses in three groups (of which half have been excavated), 25 tent rings and 30 graves. this is believed to be the locality where douglas clavering encountered the last inuit seen alive in this part of east greenland in 1823. storbukta has been used for the same feature. dødemandsbugten 74ø (74°07.3´n 20°53.2´w). name used for the station on south clavering ø at dødemandsbugten, built in 1943 as a headquarters for nordøstgrønlands slædepatrulje. it replaced the burnt out station eskimonæs, and was itself succeeded in 1944 by daneborg. it has also been known as ny station. (daumanne vågen.) dødemandsdalen 74ø (74°08.1´n 20°55.5´w). name occasionally used by danish hunters in the 1930s for skrællingedalen, a valley on south clavering ø draining into dødemandsbugten. dødemandstomten 74ø (74°07.1´n 20°55.0´w). coastal stretch of dødemandsbugten, se clavering ø. the name was used by danish hunters about 1931 because of the numerous inuit house ruins (tomt = building site). see also daudmannsøyra. dødemandstoppene 69ø-69 (69°30.0´n 29°28.0´w). mountain range on the east side of grønlands styrelse gletscher, in the high plateau region south of scoresby sund. the name originated from martin lindsay’s 1934 british trans-greenland expedition, and has appeared on maps in the forms mountains of the dead and de dødes bjerg (lindsay 1935). the mountains looked black and sinister when first seen, with a likeness to the pyramids that were the graves of the pharaohs. dødis sø 75ø (75°20.3´n 20°04.8´w). lake in hochstetter forland where samples were taken for radiocarbon age determinations (björck et al. 1994; cremer et al. 2008). dødørnryggen 72ø (72°08.0´n 24°59.6´w). name used for the sw ridge of dansketinden by the 1996 scottish mountaineering club expedition. part of the ridge was climbed, and it was described as resembling a dead eagle lying on its back. dåsen 71ø-397 (71°37.5´n 22°56.2´w). mountain 780 m high on sw wegener halvø. named by katharina perch-nielsen during the 1967–72 ggu scoresby sund expeditions for its shape (dåse = box, can). 161 e e. horse-shoe mountain 71ø (71°40.0´n 22°18.9´w). eastern of the two ridges of hesteskoen on canning land. the name is used by säve-söderbergh (1937). east cape 73ø (73°53.7´n 20°01.2´w). eastern cape of jackson ø. the name is used only in the english edition of koldewey’s 1869– 70 narrative (koldewey 1874). the expedition anchored here on 1 august 1870. east icecap 69ø (69°55.0´n 26°00.0´w). name used in a report of the 1969 watkins bjerge expedition for the present geikie plateau, an ice-covered plateau south of scoresby sund. east island – see orienteringsøerne. east plateau – see western upper terrace. east pond 72ø (72°14.4´n 23°55.0´w). name used by the 1974 joint biological expedition for a small lake near langdyssen, east of mestersvig airfield. easter glacier 72ø (72°01.0´n 24°00.0´w). name used in reports of the 1962 oxford university expedition for the present østre glet scher in the north werner bjerge. the name appears to have arisen from a mis-translation of ‘østre’ as ‘easter’ (østre = eastern). eastern circus valley 73ø (73°08.9´n 23°13.8´w). name used by gunnar säve-söderbergh during the 1931–34 treårsekspedi tionen, together with western circus valley, for two small cirqueshaped valleys on the north slopes of celsius bjerg, ymer ø (säve-söderbergh 1932). eastern mountain 71ø (71°39.5´n 22°47.0´w). mountain on eastern wegener halvø, the present tårnet. the name is found in sävesöderbergh (1937). eastern upper terrace – see western upper terrace. ebbe gletscher 76ø-340 (76°15.0´n 25°24.0´w; map 4; fig. 21). glacier in south dronning louise land flowing ne into budolfi isstrøm. named by the 1952–54 british north greenland expedition after the danish journalist and diplomat ebbe munck [1905– 74]. munck began his close association with greenland as a mem ber of the 1924 scoresbysund expedition to found the new colony, and took part in several other expeditions to east greenland. he had assisted c.j.w. simpson in the planning of the 1952–54 british north greenland expedition, and was on the expedition committee. ebeltoft vig – see æbeltoft vig. ebensbjerge 72ø (72°03.6´n 24°58.1´w; map 5). summit 2510 m high between the heads of gullygletscher and storgletscher, northern stauning alper. climbed and named by the 2007 smc east greenland expedition. the name derives from a personal name. eckhorn 71ø (72°00.4´n 25°57.4´w). mountain about 2230 m high in the northern stauning alper. named and first climbed by hans gsellman’s 1957 expedition (eck = corner). fantin (1969) and bennet (1972) give different locations for this peak. fantin locates it close to diadem and bavariaspitze, between the heads of kirk brae and storgletscher. eckspitze – see hjørnespids. eckturm 71ø (71°46.0´n 25°41.8´w; map 5). peak in the ne part of the borgbjerg gletscher region, southern stauning alper. probably named by the 1977 schwäbische stauning alper expedition. edam kulle 76ø-106 (77°00´n 18°40.0´w; map 4). highland area in east germania land. the land van edam was reported as discovered at about this latitude by a dutch whaler in 1655, and the name is found on several early dutch charts (e.g. joannes van keulen’s atlas printed in 1785). the name was adopted by the 1906–08 danmark-ekspeditionen, and placed on the most prominent point of germania land as viewed from the sea. (edams kulle.) edderfugldal 71ø-343 (71°54.8´n 22°39.3´w; map 4). valley north of fleming fjord, about 4 km west of kap biot. named during lauge koch’s 1958 expedition by k. grasmück and rudolf trüm py, for the large numbers of eider ducks observed here. edinbrae 72ø (72°04.5´n 24°26.4´w; map 5). name used by bennet (1972) for a glacier draining from the east flank of the stauning alper north wards into skeldal. edinburgh 71ø (71°43.6´n 25°14.0´w; map 5). mountain 2010 m high north of jupiter gletscher, south stauning alper. first climbed by james clarkson’s 1961 expedition, and named after edinburgh castle, scotland. edla skær 73ø (73°53.5´n 19°59.1´w). two skerries, one of them 8 m high, about 1400 m due east of the se point of jackson ø. they are said to have been found by l.m. coulet-svendsen, first mate on the gustav holm in 1930. girl’s name. the name is used in den grønlandske lods (1968). edvard bay dal 71ø-388 (71°23.0´n 27°28.0´w; maps 3, 4). major ne–sw-trending valley between martin karlsen bugt and eielson gletscher. named during the 1967–72 ggu scoresby sund expeditions after edvard bay, the geologist of carl ryder’s 1891–92 expedition. see also bay fjelde. (edward bay dal.) edvard ø 76ø-21 (76°36.0´n 21°21.0´w; map 4). island in the west part of dove bugt, so named by the 1906–08 danmark-ekspedi tionen. probably named after a member of henning bistrup’s family (j. løve, personal communication 2009). (edvards ö, edwards island.) edward bailey gletscher 71ø-420 (71°11.0´n 26°17.0´w; map 4). glacier more than 40 km long in eastern renland. named during the 1967–72 ggu scoresby sund expeditions by brian chadwick, after sir edward bailey [1881–1965], a british geologist noted especially for his work on the caledonian rocks of scotland. the rocks of the scoresby sund region form part of the circum-atlantic caledonian orogenic province. edward ensom plateau 70ø (70°40.2´n 22°43.3´w). small plateau on the west side of hurry inlet, north of moskusoksekløft. named by hermann aldinger during the 1931–34 treårsekspeditionen (al dinger 1935). egede land 69ø (c. 69°30´n 26°00´w). some atlases place this name in the region south of scoresby sund (e.g. bartholomew 1920). hans egede [1686–1758], ‘greenland’s apostle’, travelled to west greenland in 1721 in search of the remnants of the lost viking settlers, and founded a mission to serve the greenlandic inuit near present-day nuuk [godthåb]. ehrenberg fjeld 74ø-69a (74°26.5´n 21°52.0´w; map 4). mountain 1239 m high in east payer land, west of kap ehrenberg. the name came into general use among danish and norwegian hunters in the 1930s, and appears to have first been used on nsiu maps (lac mann 1937) in the form ehrenbergfjellet. eidechselspitze 72ø (72°05.7´n 25°47.9´w). snow summit about 2500 m high west of trekantgletscher, eastern nathorst land, climbed and so named by wolfgang weinzierl’s 1970 expedition. the name translates as ‘lizard peak’. eielson gletscher 71ø-67 (71°10.4´n 28°00.0´w; map 4). glacier at the head of rypefjord. this is one of the new names on the 1932 edition of the geodætisk institut 1:1 million scale map, drawn on the basis of lauge koch’s aerial observations during the first two summers of the 1931–34 treårsekspeditionen. the name comme morates carl benjamin eielson [1897–1929], a noted aviator who had made a pioneer flight with george h. wilkins in 1928 from barrow, alaska to green harbour, spitsbergen. eielson was especially noted for his ambulance flying in alaska. eiger 74ø-70 (74°26.4´n 21°00.2´w). cliff about 800 m high on the north side of clavering ø, so named by julius payer during karl koldewey’s 1869–70 expedition for its resemblance to the mountain of the same name in the bernese oberland, switzerland. (eigerfjellet.) eiger fjeld 73ø-672 (73°40´n 26°37´w). mountain about 2000 m high in central andrée land, on the north side of gneisdal. named by john haller during lauge koch’s 1949–51 expeditions, after the 162 mountain of the same name in the bernese oberland, switzerland. eigerhytta 74ø (74°26.1´n 20°56.1´w). norwegian hunting hut in the corner of lerbugt, north clavering ø, east of the mountain eiger. it was built in august 1939 as a base hut for glaciological studies by hans w:son ahlmann and kåre rodahl, and is also known as leirvågen. eigil elv 74ø-118 (74°19.7´n 21°42.1´w; map 4). large river on west clavering ø, draining west into the delta tangen. named during lauge koch’s 1930 expedition in the form eigil river, possibly after eigil riis-carstensen [1892–1953], a naval officer who was commander of the godthaab during the 1930 expedition. see also riis-carstensens dyb.(eigilelva.) eigil sø 76ø-325 (76°43.0´n 25°05.0´w; map 4; fig. 21). lake in west dronning louise land, east of revaltoppe. named by the 1952–54 british north greenland expedition after eigil knuth [1903–96], sculptor and archaeologist, noted for his numerous expeditions to greenland between 1932 and 1994 and the discovery of the early inuit independence cultures. c.j.w. simpson had first discussed his plans for an expedition to dronning louise land with eigil knuth in 1950 while in greenland, and knuth had made possible simpson’s 1951 reconnaissance of sælsøen as a means of access. (eigel sø.) eigtvedsund 75ø-34 (75°56.0´n 20°15.2´w; map 4). sound south of trums ø in the mouth of bessel fjord. named by henning bistrup during the 1906–08 danmark-ekspeditionen as eigtved sund, after his future wife ellen marie birgitte eigtved. her father was carl anton eigtved [1841–1916], a lieutenant in the danish navy (j. løve, personal communication 2009). eilan donan 72ø (72°02.3´n 25°22.2´w; map 5). rock peak about 1500 m high on the east side of dammen, north stauning alper. first climbed by malcolm slesser’s 1958 expedition, and named after eilan donan castle, scotland, a stronghold of clan macrae, now a memorial. einarfjellet 73ø (73°26.3´n 23°20.0´w). mountain about 1200 m high on the south side of gauss halvø, corresponding to part of smith woodward bjerg. the name is found on an nsiu map (1932a), and was giver for einar, one of the original norse settlers of greenland. eirik raudes land 71ø, 72ø, 73ø, 74ø, 75ø (71°30´n–75°40´n 18°–28°w). land area of east greenland claimed for norway by halvard devold when he raised the norwegian flag at myggbukta on 29 june 1931. the action was supported on 10 july 1931 by the norwegian state, and led to the court case at the international court of justice at the hague. named after eirík rauða (eric the red), a norseman banished from iceland who was noted for his discovery of greenland in 982, and its colonisation in 986. (eirik-raudes-land.) eirik raudes tinde 72ø (72°08.1´n 25°03.3´w). name occasionally used for norsketinden, a 2797 m high peak in the north stauning alper. the name occurs in accounts of the first ascent on 7 august 1954 by the dansk–norsk grønlandsekspedition. see also eirik raudes land and norsketinden. (erik rødes tinde, eirik den rødes tinde.) ejnar gletscher 76ø-333 (76°26.0´n 24°12.8´w; map 4; fig. 21). short glacier in south dronning louise land flowing west into vedel sø. named by the 1952–54 british north greenland expedition after the danish explorer ejnar mikkelsen [1880–1971]. see also ejnar mikkelsen gletscher. ejnar mikkelsen gletscher 75ø-48 (75°39.0´n 22°27.7´w; map 4; see also fig. 81). large n–s-trending glacier draining south into the head of smallefjord and east into the head of bredefjord. the name first appeared on the 1932 edition of the geodætisk institut 1:1 million scale map, and derives from lauge koch’s aerial observations during the 1931–34 treårsekspeditionen. ejnar mikkelsen [1880–1971] was one of denmark’s most eminent arctic explorers. a member of the 1898–1900 amdrup expedition, and leader of the 1909–12 alabama expedition, he was the main instigator and leader of the 1924 expedition to found scoresbysund. from 1933 to 1950 he was inspector for east greenland. ekkodal 72ø-343 (72°11.4´n 22°40.9´w). valley on se traill ø, so named during lauge koch’s 1956–58 expeditions by hans-peter heres, for an echo. ekspeditionshuset 72ø (72°07.9´n 23°51.7´w; map 5). house built on the west side of mesters vig by prospecting teams associated with lauge koch’s 1948–49 expeditions. in early 1973 it was swept away by an avalanche, and a new house, nyt ekspeditionshus, was built in 1974 about 200 m to the south. (expeditionshus.) ekstra bladets varde 72ø (72°51.6´n 26°51.7´w). cairn on the north side of dickson fjord, suess land, built on 6 august 1932 by c. eugène wegmann and aage poulsen. its name records the support given to the 1931–34 treårsekspeditionen by ekstra bladets øst grønlands fond. the cairn record was recovered by a ggu party in 1975. ekstraelv 74ø-205 (74°00.0´n 21°52.0´w). river in home forland, sw of kap stosch, so named during the 1931–34 treårsekspedi tionen by eigil nielsen. the name was originally given to the first large river south of river 1. however, in 1952 the published geodætisk institut map sheet attached the name ekstraelv to the larger river farther south (the present location), and to avoid confusion nielsen’s (1935) original ‘ekstraelv’ was renamed river zero (see discussion in teichert & kummel 1976). elefantbjerg 71ø-314 (72°00.2´n 23°40.0´w). ridge 490 m high in northern scoresby land, on the north side of kolledalen. the name was used by hans kapp during lauge koch’s 1957–58 expeditions, and given for the massive, rounded ridges, supposedly elephant-like in proportions. eleonore bugt 73ø-503 (73°26.6´n 25°22.8´w; map 4). broad bay on the east coast of andrée land between teufelsschloss and grejsdalen. named by karl koldewey’s 1869–70 expedition as eleonoren-bai. this is the only girl’s name given by the expedition, and was apparently given by koldewey himself, although there is no indication as to whom she was (j. løve, personal communication 2010). a.g. nathorst’s 1899 expedition seems to have used the name in a more restricted form than the present, for the small bay at the mouth of grejsdalen. (eleonore bay, eleonoren bay, eleonores bugt, eleonora bay, eleanore bay.) eleonore sø 73ø-415 (74°00.0´n 28°10.6´w; map 4). lake in arnold escher land. so named during lauge koch’s 1951 expedition by hans r. katz because the rocks appeared to be of the same type as those found at eleonore bugt. (eleonore-see, eleonores sø.) eleonorebukta 73ø (73°28.5´n 25°02.9´w). norwegian hunting hut in the ne part of eleonore bugt at the mouth of grejsdalen. built by arktisk næringsdrift in march 1937, it was originally known as ragnildshytta. eleonorebuktdal 73ø (73°35.5´n 26°00.0´w). name occasionally used by norwegian hunters for grejsdalen, the large valley in andrée land draining into the sea on the north side of eleonore bugt. bretz (1935) used the english variation eleonore bay valley. elephant 72ø (72°13.9´n 24°37.9´w; map 5). mountain 1830 m high at the head of tårnfjeld gletscher, north stauning alper. first climbed by the 1963 imperial college expedition, and named after the london locality, elephant and castle, originally a smithy converted to a tavern in 1760. eli knudsen øer 79ø-24 (79°22.0´n 18°43.8´w; map 4). two small islands off lambert land. the name was approved in 1958, and is attributed to the work of david malmquist during lauge koch’s expedi tions. it commemorates eli knudsen, danish hunter and member of nordøstgrønlands slædepatrulje, who was shot by german soldiers at sandodden in march 1943. dobbeltøer and tvillingøer have also been used eli knudsens hytte 72ø (72°57.5´n 24°56.5´w). hut built in 1934 at the head of a small bay on the west side of maria ø. it was repaired 163 in september 1941 by eli knudsen and hans siewers. eli knudsens varde 72ø (72°43.3´n 26°11.6´w; fig. 39). prominent cairn on the north cape of kap hedlund, built by eli knudsen on 1 august 1942 when he was stationed on ella ø. eli knudsens vig 72ø (72°57.1´n 24°57.2´w). this name was reported by olsen (1965) as used for the small bay on western maria ø where eli knudsens hytte was built in 1934. elis bjerg 70ø-123 (70°55.1´n 22°41.7´w). mountain 540 m high nw of the head of hurry inlet. named by alfred rosenkrantz during lauge koch’s 1926–27 expeditions after his greenlandic assistant, eli napartok, and used in the form eli fjæld and mount eli fjæld. a.g. nathorst’s ammonitbjerg had been erroneously placed at this location, but was moved by rosenkrantz (1934) further inland. elisabeth bjerg 73ø-530 (73°42.4´n 25°11.4´w). mountain 588 m high on the west side of geologfjord, andrée land, named by a.g. nathorst’s 1899 expedition as elisabeths berg, probably after nathorst’s daughter elisabeth. see also kap elisabeth. elisabethsminde 72ø-252 (72°10.9´n 24°48.9´w; map 5). moun tain 2260 m high in the northern stauning alper, climbed by peter braun and fritz schwarzenbach in august 1951 during lauge koch’s 1951 expedition. the original name was elisabethstinde (braun 1953), and was given for elisabeth fränkl-fischer, wife of erdhart fränkl. the significance of the change from ‘tinde’ to ‘minde’ (= memorial) is uncertain, as elisabeth was still alive in 1990. braun and schwarzenbach (whose wife is also named elisa beth) had assisted fränkl in his geological work in this region in 1950–51. elizabeth sharon sø 76ø (76°06.5´n 20°14.7´w). name used on 1952 wac maps for the present gunner andersen sø in eastern ad. s. jensen land. ella ø 72ø-49 (72°51.0´n 25°03.0´w; maps 3, 4; figs 8, 29). large island at the mouth of kempe fjord, inner kong oscar fjord. named by a.g. nathorst during his 1899 expedition as ellas ö, after his wife amy rafaela (ella) windahl [1858–1936]. (ella is land, ella-öya, ellainsel.) ella ø station 72ø (72°52.6´n 25°06.7´w; fig. 40). name often used for lauge koch’s scientific station built in august 1931 at the head of solitærbugt, an excellent harbour in northern ella ø. the main house, often referred to as ørnereden, was continuously manned from 1931 to 1943 and 1947 to 1952, and used by lauge koch’s expeditions as a summer station until 1958. during the war years the station was taken over by nordøstgrønlands slæde patrulje and was occasionally known under the code name bluie east 4. the station was damaged and the radio masts cut down by german soldiers in 1943. a variety of additional houses were built later, of which a canteen and barrack east of the main house still stand. eight small houses were built in 1950–51 to house 51 staff, including 18 aircrew, engaged in aerial photography with three catalina flying boats. sirius make use of the last of these houses, largely rebuilt, and have added several storehouses for their own use. (ellaøstation.) ellemandsbjerge 72ø-85 (72°27.2´n 22°10.3´w; map 4). mountain range on the north side of mountnorris fjord, eastern traill ø. named during the 1931–34 treårsekspeditionen by ove simonsen after the danish locality of the same name at helgenæs, jylland. elsa dal 73ø-116 (73°23.5´n 23°08.4´w). ravine on the sw coast of gauss halvø. so named by gunnar säve-söderbergh during the 1931–34 treårsekspeditionen, after elsa warburg, an assistant professor at the university of uppsala, sweden. elsinore fjeld 72ø-486 (72°13.2´n 24°34.5´w; map 5). mountain 1829 m high north of bersærkerbræ, north stauning alper. first climbed by john hunt’s 1960 expedition, and named elsinore after the castle in the danish town of helsingør, made famous by shakespeare’s ‘hamlet’. the second ascent of the peak was by the 1963 imperial college expedition. elvdal 70ø-308 (70°29.3´n 21°57.1´w). valley adjacent to the settlement of scoresbysund, south liverpool land (elv = river, dal = valley). it was so named during the 1924 colonisation expedition (pedersen 1926). elvdalen 70ø-63 (70°28.2´n 26°16.0´w). nw–se-trending valley north of hekla havn on danmark ø, which carries the only significant stream on the island. so named by carl ryder’s 1891–92 expedition. elveidet 72ø (72°17.5´n 24°08.9´w). name used by the møre expedition for a norwegian hunting hut built in august 1930 east of the mouth of skeldal, northern scoresby land (rogne 1981). it was named for the low ‘eid’ (= meadow) between the river and the hut. the hut is now usually known as skeldal-hytte and has also been called havnhytta. (elv-eidet.) elvsborg 74ø-100 (74°07.9´n 20°39.9´w). norwegian hunting station on se clavering ø, se of brinkley bjerg. it was named after the large river (= elv) beside the station. built in 1927 by the 1927–29 hird expedition, it has also been referred to as hirds fox farm. trapped foxes were kept alive in cages until the condition of the fur was optimal, and were then killed. (elfsborghytten, elfsborg.) emilia bjerg 72ø-421 (72°53.1´n 27°26.7´w; map 4). mountain about 1700 m high in sw suess land. named during the 1931–34 treårsekspeditionen by eugène wegmann as mt. emilia, allegedly after an italian locality. fig. 39. prominent cairn on the north cape of kap hedlund, built by eli knudsen on 1 august 1942, and generally known as eli knudsens varde. the seated geologists are reading a copy of eli knudsen’s cairn record. 164 emmabjerg 71ø-298 (71°58.0´n 26°04.3´w). mountain 2540 m high on the south side of furesø, nathorst land, named by hans zweifel during lauge koch’s 1954–55 expeditions. girl’s name. published geodætisk institut map sheets show the location several kilometres further to the south. emmanuel fjeld 71ø-361 (71°57.8´n 25°06.5´w; map 5). peak about 2400 m high in the stauning alper. climbed by a cambridge university expedition on 3 august 1963, and named after emma nuel college, cambridge, founded in 1584. (emmanuel.) emmanuel gletscher 70ø-243 (70°51.1´n 21°49.8´w; map 4). glacier in liverpool land draining eastwards to reach the sea near janus ø. named by brian roberts after emmanuel college, cam bridge, which had given financial support to his 1933 cambridge expedition. see also emmanuel fjeld. endalip kangersiva 70ø (70°28.4´n 21°54.5´w). name recorded by the local scoresbysund newspaper in 1984 as in use for the bay close to scoresbysund, south liverpool land, officially known as ittoqqortoormiit kimmut kangertivat [amdrup havn]. endeløs 73ø-377 (73°42.7´n 25°33.5´w; map 4). long glacier in ne andrée land draining east via morænedal to geologfjord (endeløs = endless). named during lauge koch’s 1948–50 expeditions by erdhart fränkl, who traversed the glacier many times with his assistant during geological field work, transporting equipment and food. (endeløs gletscher.) eneboerdal 72ø-394 (72°02.8´n 23°09.3´w). valley on the nw side of antarctic havn, north scoresby land. the name was used by hans kapp during lauge koch’s 1957–58 expeditions (eneboer = hermit). enehøj 75ø-69 (75°06.4´n 18°27.3´w). solitary hill 82 m high on central shannon. the name is attributed to the wintering party at kulhus during the 1931–34 treårsekspeditionen. named for its isolated location, and possibly also after the island of this name in nakskov fjord, denmark, which was owned by peter freuchen from 1926 to 1940. engdalen 73ø-599 (73°13.1´n 27°16.9´w; map 4). valley in south frænkel land, named by gunnar thorson during the 1931–34 treårsekspeditionen because of the rich vegetation (eng = meadow). engelhardts sund 76ø (76°18.8´n 20°40.0´w). sound between na nok ø and tvillingerne, sw dove bugt, corresponding to the pres ent jægersund. the name appears in the account of the 1932 fig. 40. lauge koch’s expedition base on ella ø, ella ø station, viewed from the east. one of the expedition’s norseman aircraft is moored close to the shore of solitærbugt in the foreground. the large house with the white-painted gables at centre is the original expedition house built in 1931, known as ørnereden. the overwintering geologists and telegraphists lived in this house during the winter. the two houses at the right near the beach are a kitchen house and a barrack. on the beach is a large depot of fuel drums. the houses at the left were built in 1950–51 to house the crews of the catalina aircraft and personnel engaged with geodætisk institut aerial photography. the john haller photograph collection, geus archive. 165 gefion expedition (jennov 1935), and was given for svend engel hardt, a lawyer who was one of the founders of the nanok hunting company. (engelhards sund.) engelsborg 70ø (70°17.8´n 24°44.2´w). cliff about 1000 m high on volquaart boon kyst, west of solgletscher, the present stejlfjeld. the name was reported by henning bistrup in 1939 as communicated to him in 1930 by johan petersen (governor of scoresbysund) aboard the gustav holm (engelsborg = angels castle). engledal 72ø-346 (72°13.3´n 22°35.4´w). valley on se traill ø, so named during lauge koch’s 1956–58 expeditions by hans peter heres (engle = angels). engpasset 74ø (74°24.2´n 20°01.9´w). pass in wollaston forland east of kuppelpasset, between summits 450 m and 703 m (eng = meadow). the name was used by andreas vischer during 1937 field work with lauge koch’s expeditions (vischer in: koch 1955). enhjørningen 70ø-260 (70°09.9´n 24°02.6´w). prominent peak 1730 m high on volquaart boon kyst, so named during the 1931–34 treårsekspeditionen by laurits bruhn for its single spire (enhjørning = unicorn). enhjørningen 80ø-79 (80°12.1´n 20°53.4´w). mountain in southern kronprins christian land, south of sæfaxi elv. named during lauge koch’s 1952–53 expedition by erdhart fränkl for its single summit peak. enhjørningen dal 71ø-176 (71°34.5´n 23°10.7´w; map 4). valley draining from jens munk plateau ne to fleming fjord. the name was one of a group of names given by the place name committee in 1939, and commemorates the enhjørningen, one of jens munk’s ships used on his voyage in search of the nw passage in 1619. eremitdal 73ø-603 (73°49.5´n 26°00.0´w; map 4). major e–wtrending valley in north andrée land draining into geologfjord. the name was first used in botanical reports of the 1931–34 treårsekspeditionen (gelting 1934), and may record the find of a solitary inuit house ruin at the mouth of the valley in 1931 (eremit = hermit). eremitdalshytten 73ø (73°48.9´n 25°36.3´w). norwegian hunting hut built for arktisk næringsdrift, probably in 1936, on the north side of the mouth of eremitdal, ne andrée land. it is also known as wintherheimen. (eremitdalhytten.) eremitten 74ø-195 (74°42.3´n 23°21.4´w). nunatak north of wordie gletscher, discovered on a journey along the inland ice margin in 1932 during the 1931–34 treårsekspeditionen. so named by th. johansen and curt teichert because of its isolation (eremit = hermit). (eremit nunatak, mt. eremit.) erik rødes tinde – see eirik raudes tinde. erik s. henius land 81ø-70 (81°30.0´n 11°48.4´w; maps 1, 4). coastal area between nordostrundingen and nakkehoved, ne kronprins christian land. so named by the 1906–08 danmarkekspeditionen after erik semmy henius [1863–1926], a danish businessman and consul, generous supporter of danish arctic exploration and a member of the 1906–08 danmark-ekspedi tionen committee. eros 71ø (71°37.8´n 25°14.1´w; map 5). mountain about 2018 m high at the head of mercurius gletscher and oxford gletscher, south stauning alper. it was first climbed by the 1975 scottish scoresby land expedition, and named after the minor planet eros which has a highly eliptical orbit; the planet was named after the goddess of love. erratic bloc bay 73ø (73°15.1´n 22°12.2´w). name used by bütler (1954) for a small bay at kap franklin, gauss halvø, where erratic boulders were found. this name is only used on bütler’s maps, and in the text of his report is replaced by findlingsbucht. erste weisse – see første hvide. erzberg 71ø (71°59.2´n 24°15.3´w). name commonly found in swedish and german publications (erz = ore) describing the molyb denum deposit at malmbjerg, werner bjerge (e.g. sjögren 1962). see also malmbjerg. escher von der linth gletscher 71ø (71°57.8´n 23°45.6´w). name occasionally used for the present cirkusgletscher which drains from the eastern werner bjerge to blomsterdal. the name is found in the description by brooks et al. (1982) of samples collected by peter bearth in 1953–54. see also arnold escher land. eskdal 72ø-517 (72°40.4´n 23°47.7´w; map 4). valley on sw traill ø draining south into karupelv. named by geoffrey halliday following his botanical work during the 1961 leicester university expedition and the 1971 northern universities expedition, probably after eskdale in the lake district of northern england. eskimo land 74ø (74°32.1´n 18°50.1´w). name used by dahl (1925) for the peninsula west of germaniahavn, sabine ø, where kolde wey’s observatory was built in 1869–70. it was given for the presence of inuit (eskimo) ruins. eskimobugt 71ø-355 (71°38.6´n 27°11.9´w). bay on the north side of inner nordvestfjord, southern nathorst land, on the shore of which are well-preserved inuit ruins. the site has been known since the 1930s when visited by eduard wenk, but was named by the 1963 geodætisk institut expedition who noted the ruins while anchored here with the tycho brahe. the published geodætisk institut 1:250 000 scale map sheet (71 ø.2) locates the bay incorrectly about 8 km further north. eskimohamna 74ø (74°05.6´n 21°16.0´w). name used on the nsiu (1932a) map for østhavn in south clavering ø, beside eskimonæs scientific station. eskimonæs 74ø-126 (74°05.5´n 21°17.2´w; maps 2, 4). prominent peninsula on south clavering ø, named by lauge koch’s 1929–30 expeditions as eskimonæsset, for the inuit (eskimo) settlement of four houses here, of which two were excavated (glob 1946). the same name is now officially used for the ruins of the scientific station built by koch in 1931 in the bay ne of the cape (74°05.7´n 21°16.8´w). eskimonæs station was used as a wintering station by scientists from 1931 to 1939, and from 1941 to 1943 was the headquarters of nordøstgrønlands slædepatrulje. the main building was damaged by a german patrol on 25 march 1943, and the site bombed by the us air force on 14 may the same year. the burntout remains of the station are a conspicuous memorial to war-time activities, and remain essentially undisturbed. the names south cape and foxtrap point have also been used for the peninsula. eskimonæs 80ø-7 (80°25.9´n 15°46.3´w; maps 1, 4). peninsula on the ne coast of holm land. so named by the 1906–08 danmarkekspeditionen as eskimonæsset, because ruins of a large inuit settlement were found here in march 1907. the same name is used for the sirius hut at the cape slightly farther north. (eskimo naze, eskimo peninsula, esquimo peninsula.) eskimovig 74ø-91 (74°05.7´n 21°07.5´w). small bay on the south side of clavering ø, named by j.m. wordie’s 1926 expedition as eskimo bay because of the many inuit (eskimo) ruins. a settlement of 25 winter houses occurs here, of which 18 have been excavated (glob 1946). on norwegian maps (lacmann 1937) this site is referred to as breivika or breidvik, their eskimohamna (the present østhavn) lying farther west. (eskimobugt.) essemmceebrae 71ø (71°59.7´n 25°14.7´w; map 5). minor glacier on the south side of sefström gletscher, stauning alper. explored by the 1998 scottish mountaineering club expedition, and so named after the club (smc). etagefjældet 74ø (74°16.1´n 19°42.3´w). this name has been used by danish hunters of østgrønlandske fangstkompagni for a mountain in south wollaston forland, probably the present herschell bjerg. it probably derives from the stratified appearance of the basalt rock formations (etage = tier, floor). etzelbreen 74ø (74°21.9´n 21°16.9´w). glacier draining nw in north clavering ø. so named on the nsiu maps of lacmann (1937) after etzel (attila), king of the huns, the second husband of kriemhild in the german epic poem from c. 1200, the nibe 166 lungenlied. eugen-heinz tinde 71ø (71°47.4´n 25°37.8´w; map 5). peak 2415 m high in the ne part of the borgbjerg gletscher region, southern stauning alper. probably first climbed and named by the 1977 schwäbische stauning alper expedition. eva ø 79ø (79°18.9´n 18°56.3´w). small island off ne lambert land. the name was used by the 1996 mylius-erichsen’s minde ekspedition. evald gletscher 77ø-59 (77°16.8´n 20°13.5´w; map 4). glacier in ne søndermarken, on the south side of c.f. mourier fjord. so named by david malmquist during the 1931–34 treårsekspedi tionen after evald hellman, an old friend and class-mate, later chemist and assistant at the chemical institute in uppsala, sweden. evans bjerg 73ø-286 (73°21.0´n 22°48.9´w). mountain on the sw coast of gauss halvø. named during the 1931–34 treårsekspedi tionen by gunnar säve-söderbergh as mt. evans, after john wil liam evans [1857–1930], a widely travelled british geologist, who was an authority on the devonian deposits of great britain. eventyrfjelde 76ø-344 (76°05.5´n 24°22.4´w; maps 2, 4; fig. 21). nunatak with summits reaching 2000 m south of a.b. drachman gletscher, south dronning louise land. the name was given by the 1952–54 british north greenland expedition, and arose apparently because this outlying area was hardly investigated by the expedition. anything they could say about it was likely to be a fairy tale (= eventyr). evers gletscher 73ø-593 (73°41.5´n 29°25.0´w; map 4). glacier between hvidbjørn nunatakker and knud ringnes nunatak. the area was first explored by arne høygaard and martin mehren in 1931, and in 1932 overflown by lauge koch during the 1931–34 treårsekspeditionen. the name first appears on the 1932 geo dætiske institut 1:1 million scale map, and was given for the captain of the hvidbjørn in 1932, who had assisted koch. this was probably christian vilhelm evers [1887–1966], who served in the danish navy from 1908 to 1945. ewaldfjellet 74ø (74°24.3´n 21°09.9´w). mountain 1500 m high in north clavering ø. so named on the nsiu maps of lacmann (1937), after erich ewald [b. 1884], a german minister who encouraged the nsiu work on maps of east greenland which were drawn in berlin. (ewald-fjellet.) expeditionshus – see ekspeditionshuset. eyvind fjeld gletscher 74ø-176 (74°07.9´n 27°00.0´w; map 4). tributary glacier on the north side of adolf hoel gletscher. named by arne høygaard and martin mehren in 1931 after their fellow student eyvind fjeld, the initiator of their expedition who had been unable to join them on their crossing of the inland ice. f f. graae gletscher 72ø-414 (72°06.8´n 28°42.3´n; map 4). glacier at the head of nordvestfjord, on the ne side of charcot land. the name first appeared on the 1932 1:1 million scale geodætisk institut map prepared by lauge koch during the 1931–34 treårs eks pedi tionen. it was named after frederik graae [1875–1948], under-secretary of state and vice-president of the treårs ekspe ditionen committee. graae had been particularly helpful in obtaining support for koch’s 1929 expedition (koch 1930b). the name appears incorrectly on some maps as graah gletscher (e.g. koch & haller 1971). f. toula plateau 77ø-145 (77°05.0´n 18°46.4´w; map 4). plateau in germania land, se of fladebugt. named during lauge koch’s 1956–58 expeditions by john haller after the austrian geologist franz toula [1845–1920], who had worked up collections made in this area by karl koldewey’s 1869–70 expedition. fakirgryde 72ø-385 (72°03.9´n 23°24.3´w). glacier-filled, basinshaped valley north of the head of segldal, northern scores by land. named by hans kapp during lauge koch’s 1957–58 expeditions. falkberget 74ø (74°33.9´n 19°18.2´w). norwegian hunting hut se of falkebjerg, on the north side of falskebugt, wollaston forland, built by the hird expedition in august 1928. the name of both hut and mountain derive from the gyrfalcon, formerly common in the region. the hut has also been known as taymors fjell and falske bugt hytten. (falkeberget, falkenberg, falkerbjerghytte.) falkebjerg 74ø-287 (74°34.4´n 19°19.6´w). mountain 307 m high in ne wollaston forland, north of falskebugt. see also falkberget. falkeelv 70ø-111 (70°52.3´n 22°53.1´w). river draining into ugle elv on the west side of the head of hurry inlet. named by alfred rosenkrantz during lauge koch’s 1926–27 expeditions as falkon river. roberts (1935) placed the name in error on a south-flowing river west of nathorst fjeld. (falke elv.) falkepynt 70ø-68 (70°27.7´n 26°29.7´w; map 4). peninsula in ne gåseland. named by carl ryder’s 1891–92 expedition as falke pt. gyrfalcons were seen by the expedition on several occasions. falkeryg 74ø-221 (74°00.9´n 21°33.1´w). minor ridge between river 9 and river 10, on the north flank of frebold bjerg, home forland. so named by eigil nielsen during the 1931–34 treårseks pedi tionen, after the gyrfalcon. falketind 73ø (73°07.9´n 23°07.8´w). eastern peak of celsius bjerg, about 880 m high, nw of kap humboldt on ymer ø. so named on an nsiu map (1932a). falkonerklippe 76ø-135 (76°28.1´n 26°26.2´w; map 4). nunatak in sw dronning louise land, west of helgoland. so named during j.p. koch’s 1912–13 expedition because two gyrfalcons were seen here on 5 may 1913. (falkonerklippen, falkoner klippen.) false col 72ø (72°08.4´n 24°54.0´w; map 5). col between the heads of bersærkerbræ and skjoldungebræ, between bosigran and ken sington, stauning alper. named by the queen mary college expedition which climbed the col from the bersærkerbræ side on 15–16 july 1968. falske bugt hytten 74ø (74°33.9´n 19°18.2´w). norwegian hunting hut built in august 1928 on the peninsula north of falskebugt, wollaston forland, by the hird expedition. it was originally known as falkberget, and has also been known as taymors fjell. see also falskebugt. falskebugt 74ø-55 (74°33.3´n 19°21.5´w). bay in ne wollaston forland. so named by karl koldewey’s 1869–70 expedition as falsche bai, because the bay appears from the sea to be much larger than it is, due to the low ground at its head (falske = false). (falschen bai, falschen bay, falsche bay, flache bugt.) falskenæs 74ø-313 (74°06.1´n 21°19.3´w). prominent peninsula nw of vesthavn, near eskimonæs, on the south coast of clavering ø, similar to and sometimes mistaken for the peninsula eski monæs. named by the wintering party at the station during the 1931–34 treårsekspeditionen. (falske næs.) falsterselv 70ø-93 (70°51.1´n 24°00.0´w). river in jameson land flowing west into hall bredning. named during the 1931–34 treårsekspeditionen by laurits bruhn after the island of falster, denmark. fame øer 70ø-158 (70°48.9´n 22°29.2´w; map 4). group of two large and three small islands at the head of hurry inlet. named by william scoresby jr. during his 1822 expedition as the fame islands, after his father’s ship the fame of hull, the first to explore hurry inlet. the fame was a teak-built ship, a prize from the french purchased by william scoresby sr. in 1817. he sailed it to the whale fishery from 1819 to 1822, and retired after the fame was destroyed by fire at stromness in the orkneys in 1823. (fame öar, fame öarne, îles fame.) fangergletscher 73ø-388 (73°36.7´n 25°53.3´w). glacier in an drée land draining south into grejsdalen. named during lauge koch’s 1948–50 expeditions by erdhart fränkl (fanger = hunter, sealer). (fanger gletscher.) 167 fangersund 76ø-283 (76°16.2´n 21°25.9´w). sound inside the skerries off the coast of ad. s. jensen land between soranerbræen and syttendemajfjorden. named by the 1938–39 mørkefjord expedition for the danish hunters who operated in the region. fangsthytte, fangsthus, fangststation – these names are in general use on official topographic maps for danish or norwegian hunting huts and stations, most of which have individual names given in this catalogue, and which are described in detail by p.s. mikkelsen (1994, 2008). the representation of a hut or station on a published map is no guarantee of its continued existence, as many have been destroyed by katabatic winds or by bears, and only a few of the larger stations are now maintained by sirius. fangsthytten 75ø-99 (75°21.2´n 21°19.4´w). official name for the danish hunting hut on the south side of ardencaple fjord at the mouth of femdalen, built by nanok in september 1930. now a ruin (1988). this hut was usually known as femdalhytten. fangsthyttegletscher 72ø-313 (72°15.0´n 25°12.4´w; maps 4, 5). glacier in the nw stauning alper draining into alpefjord. named by john haller during lauge koch’s 1954 expedition, for the hunting hut west of the front of the glacier. faraway how 74ø-87 (74°24.2´n 23°29.9´w; map 4). nunatak about 1500 m high in the upper part of wordie gletscher, named by j.m. wordie’s 1926 expedition. it is a whimsical name for a very distant object. farimagdalen 76ø-128b (76°34.0´n 24°27.3´w; map 4). valley south of borgjøkel in central dronning louise land containing farigmagsø. named by j.p. koch’s 1912–13 expedition as far imags dalen or farimag-tal because it was easy sledging (far i mag = travel at leisure). farimagsø 76ø-128 (76°36.0´n 24°27.3´w; map 4). lake in fari magdalen, on the south side of borgjøkel, dronning louise land. named by j.p. koch’s 1912–13 expedition as farimagsvandet or farimagsøen, because it was covered by snow-free smooth ice. (farimagsee.) farsund 76ø-30 (76°51.6´n 19°34.8´w). sound between winge kyst and nørre orienteringsø, off the south coast of germania land. so named by christian b. thostrup during the 1906–08 danmark-ekspeditionen because they always travelled this way during their journeys in dove bugt; however, in his diary thostrup records that he always thought of his father (= far; thostrup 2007). (fairway, far sound.) farvefjeldet 77ø-91 (77°05.1´n 21°39.4´w; map 4). mountain on the north side of western sælsøen. named by the 1938–39 mørke fjord expedition, for the contrasting colours of the rocks (farve = colour). (farvefjældet.) farvel nunatak 77ø-140 (77°10.3´n 26°12.5´w; map 4). this ‘nunatak’ was originally two small nunataks close together, the most westerly nunataks of dronning louise land. since 1954 the melting of the ice has revealed a group of six closely spaced nunataks. so named by the 1952–54 british north greenland expedition because they were the last landmark of the expedition groups proceeding west to the ‘northice’ station on the inland ice (farvel = goodbye). fata morgana landet 79ø, 80ø (c. 80°00´n 10°00´w). elusive land area or island supposedly lying between ne greenland and spitsbergen. eigil knuth (1940) reported it was first seen by johan peter koch and aage bertelsen in 1907, and subsequently noted by lauge koch from the air in 1933, and by peter freuchen in 1935. another supposed sighting by ivan papanin during his ice-flow drift in 1937 led directly to lauge koch’s 1938 seaplane expedition from spitsbergen, which found no trace of it (koch 1940). the reported sightings were possibly of tobias ø, an island somewhat further south and 70 km from the greenland coast whose position was determined in 1993. the ‘fata morgana’ or ‘castles in the air’ are a spectacular form of mirage (see also fig. 53), common in the arctic, named after morgan le fay, king arthur’s fairy half-sister, who according to the arthurian legend lived in a crystal castle under the sea. (fata morgana øerne.) fault valley 73ø (73°09.9´n 23°21.1´w). name used by gunnar säve-söderbergh during the 1931–34 treårsekspeditionen for a deep and narrow valley on the nw side of celsius bjerg, ymer ø. in säve-söderbergh (1933) northern fault valley is used for this feature and together with southern fault valley, forms the valley known to norwegians as forkastningsdalen. faustsøen 73ø-668 (73°45.0´n 26°38.6´w; map 4). long lake in innermost eremitdal, central andrée land. named by john haller during lauge koch’s 1949–51 expeditions, for the magical, grave and mysterious scenery surrounding the lake. faust, or dr. faustus, was the hero of the legend of a german astrologer who sold his soul to the devil. favoritdal 74ø (74°28.1´n 20°39.2´w). valley on the se slope of zackenberg. the name is used as a reference locality by scientists visiting zackenberg forskingsstation (e.g. meltofte & thing 1996). (favorite valley.) favre bjerg 73ø-314 (73°56.3´n 23°17.7´w; map 4). mountain about 1900 m high in central hudson land. named by heinrich bütler during lauge koch’s 1936–38 expeditions after jean alphonse favre [1815–1890], a swiss structural geologist and pioneer of alpine geology, who became director of the schwei zerische geologische kommission. faxa sø 70ø-384 (70°13.6´n 28°44.4´w; map 4). large lake in western gåseland, draining via hjørnedal to fønfjord. it was reported by eduard wenk to have been so called throughout the 1958 summer by lauge koch, expedition members, and the crew of the flugfélag islands catalina which landed wenk and his assistants on the lake. the name is a corruption of ‘sæfaxi’, the name of the catalina. see also sæfaxi elv. (faxa-sø.) fegin elv 71ø-194 (71°12.6´n 23°51.5´w; map 4). river in jameson land draining sw to hall bredning. one of a group of names given by the place name committee in 1939, it was given for finn fegin, son of olav den hellige, who was lost with his ship in greenland in about 1028. see also lodin elv. fellenberg gletscher 73ø-330 (73°57.5´n 22°36.1´w). glacier in hudson land. named during lauge koch’s 1936–38 expeditions by wolf maync and andreas vischer after one of the pioneers of swiss geology, e. von fellenberg. lacmann’s (1937) maps use höy gaardbreen. felspingo 71ø (71°46.3´n 23°36.6´w). name used by müller (1959) for the remains of a pingo in pingo dal, north jameson land, which consists of a barrier of rock debris 29 m high across the valley (fels = rock). femdalen 75ø-35 (75°20.4´n 21°28.8´w; map 4). valley on the sw side of ardencaple fjord, named by the 1906–08 danmarkekspeditionen in the form fem-dalene because it branches into five valleys. (femdalene.) femdalen 75ø (c. 75°22´n 21°21´w). norwegian hunting hut built by arktisk næringsdrift in november 1949 on the north side of the mouth of femdalen. it had disappeared by november 1952 (p.s. mikkelsen 1994). femdalhytten 75ø (75°21.2´n 21°19.4´w). danish hunting hut on the south side of ardencaple fjord at the mouth of femdalen, whose official name is fangsthytten. the hut is said to have been made on sabine ø and transported here in september 1930. hansen (1939) described it as a miserable hut made out of packing cases. it was a ruin in 1988. (femdalshytten.) fermi klippe 76ø-307 (77°00´n 25°14´w; map 4). cliff on the north side of admiralty gletscher, nw dronning louise land. one of the names given by the 1952–54 british north greenland expedition for notable scientists, it commemorates the italian physicist enrico fermi [1901–54], considered to be one of the chief architects of the nuclear age. ferskesø 70ø-62 (70°28.8´n 26°18.2´w). lake on danmark ø, 168 draining through elvdalen. named by carl ryder’s 1891–92 expedition as ferske sø because it was the source of drinking water for the winter quarters at hekla havn. ferslew pynt [palasip qammavaajua] 70ø-305 (70°29.3´n 21°58.6´w). peninsula in rosenvinge bugt, south liverpool land, ad jacent to the settlement of scoresbysund established in 1925. named during the colonisation expedition in 1924 (e. mikkelsen 1925) after valdemar galster, owner of the ferslew press, who had done much to raise funds for the enterprise. (ferslew cape, ferslews pt.) festningen 72ø (72°42.4´n 26°47.6´w). norwegian hunting hut built for arktisk næringsdrift on the east side of stromnæs, glet scher land, in july 1934. it is also known as strømnæshytten and röhss fjord hytten. filosofbjerg 72ø-470 (72°02.0´n 26°28.9´w; map 4). mountain 1965 m high at the west end of furesø, nathorst land. named by hans zweifel during lauge koch’s 1954–55 expeditions, perhaps for its appearance, or for zweifel’s meditations while camped near the mountain (filosof = philosopher). fimbulbreen 71ø (71°54.5´n 25°08.4´w; map 5). name given to a northern branch of roslin gletscher by the 1996 norwegian stauning alper expedition. it was named after the ‘fimbulsvinter’ of nordic mythology. see fimbulfjeld. fimbulfjeld 72ø-129 (72°53.6´n 24°58.9´w). mountain 634 m high on northern ella ø, whose north side is in shadow most of the year. so named by the ella ø wintering party during the 1931–34 treårsekspeditionen, the name deriving from the ‘fimbulsvinter’ of nordic mythology, according to which three successive winters killed everything living and caused ‘ragnarok’. the wintering par ties had experienced three long, dark winters in succession, with intervening poor summers. (fimbul mt.) fimbulpasset 71ø (71°55.6´n 25°08.4´w; map 5). pass between the head of canta bræ and fimbulbreen. the name was used by the 1996 norwegian stauning alper expedition, probably the first to cross it and also responsible for naming fimbulbreen. (fimbul passet.) findelen sø [issø] 72ø-457 (72°47.1´n 28°10.0´w). lake to the north of hisinger gletscher, goodenough land, so named during the 1931–34 treårsekspeditionen by eugène wegmann, after the valley and glacier of the same name near zermatt, switzerland. (findelensee.) findlingsbucht 73ø (73°15.1´n 22°12.2´w). name used by bütler (1954) for a small bay at kap franklin. on the plates of his publication bütler uses erratic bloc bay for the same feature. both names refer to finds of fossils in ice-transported boulders in the bay. fingerbøllet 71ø-278 (71°54.7´n 24°00.2´w; map 5). mountain in the werner bjerge between sirius gletscher and aldebaren glet scher, named by peter bearth and eduard wenk during lauge koch’s 1953–54 expeditions (fingerbøllet = thimble). it was climbed by bearth’s party in 1954, and named for the shape. (fingerhut.) fingerfjeld 74ø (74°15.3´n 20°31.6´w). this name is used in den grønlandske lods (1968) for an 800 m high mountain on ne clavering ø, the present langelinie. possibly named for a fingerlike shape. fingerklippen 74ø (74°11.3´n 20°07.1´w). cape on ne clavering ø south of kap breusing, the present kap arnakke. the name is used in den grønlandske lods (1968). possibly named for a finger-like appearance. finnsbúdir 71ø, 72ø (72°01.0´n 23°08.0´w). bay mentioned in the icelandic sagas (ivar baardsson’s annaler), which according to tornøe (1944) might correspond to the present antarctic havn on the south side of kong oscar fjord. in support of his speculations tornøe noted the remains of a wooden house of supposed norse origin on the east side of antarctic havn (lavenesset), found by jonas karlsbak in 1930. the name derives from finnr (finn) fegin, a son of olav den hellige, who with his ship’s crew was said to have drowned in the harbour in 1028. see also fegin elv. finnvatnet 72ø (72°53.1´n 22°04.8´w). lake on eastern geo graph ical society ø. used only on the nsiu maps of lacmann (1937), the name was given for finn devold [1901–77], who led nor wegian hunting expeditions to east and southern east greenland in 1928–30 and 1931–33. from 1938 until his retirement he worked with the norwegian fiskeridirektorat (ministry of fisheries). finsch øer 73ø-21 74ø-266a (74°01.5´n 20°54.9´w; maps 2, 4). group of five large islands south of clavering ø, including store finsch, stille ø, kalven and lille finsch. the islands were named by karl koldewey’s 1869–70 expedition as finsch inseln, after otto friedrich hermann finsch [1839–1917], a german zoologist and ethnologist. he contributed the ornithology chapter to koldewey’s expedition narrative. (finsch islands, finschøya.) fireogtyvekilometernæsset – see fyrretyvekilometernæsset. fireskæret 76ø-279 (76°26´n 20°43´w). skerry east of godfred hansen ø in the sw part of dove bugt. named by the 1938–39 mørkefjord expedition. firkanten 73ø-220 (73°21.3´n 22°26.5´w). mountain 970 m high north of margrethedal in se gauss halvø, so named on an nsiu map (1932a) for its shape (firkant = square). firkanten 73ø (c. 73°15´n 22°20´w). norwegian hunting hut built by arktisk næringsdrift in the summer of 1937 about 5 km west of kap franklin. it was also known as funkis. no trace of it now remains. firmannsdalen 73ø (73°07.3´n 24°36.1´w). name given in 1934 to fladedal, central ymer ø, to record a successful hunting trip in the valley by four norwegian hunters (firmann = four men). the name has also been used for the norwegian hut built in 1934 at the mouth of the valley (see namdalsstua). firndalen 80ø-38 (80°20.0´n 18°00.0´w; map 4). n–s-trending valley crossing holm land. so named by eigil nielsen during the 1938–39 mørkefjord expedition because of the large glaciers (= firn) on both sides. first point of aries 71ø (71°37.5´n 25°06.0´w). peak 1944 m high on the south side of mercurius gletscher, south stauning alper. first climbed and so named by james clarkson’s 1961 expedition. the ‘first point of aries’ is the intersection of the celestial equator and the apparent annual pathway of the sun, from which celestial longitude and latitude is measured. (aries.) fiskeelv 74ø-104 (74°08.1´n 20°40.3´w). small river on se clav ering ø, named by lauge koch’s 1929–30 expeditions in the form fish river because of finds of fossil fish. fiskeplateau 74ø-224 (74°01.5´n 21°35.3´w). minor plateau be tween river 7 and river 8, on the north slope of frebold bjerg, home forland. so named by eigil nielsen during the 1931–34 treårsekspeditionen for finds of fossil fish. (fiskeplateauet.) fiskergrav 71ø (c. 71°24´n 24°35´w). locality in sw jameson land between schuchert flod and gurreholm dal where stemmerik et al. (2001) located the permian–triassic boundary on the basis of palynological data. fiskerhytten 74ø (74°27.9´n 20°39.1´w). norwegian hut built for fishing by hermann andresen’s expedition about 500 m west of zackenberg hunting station. it is also known as laksehytten. fiskerhytten 76ø (76°11.1´n 20°43.3´w). danish hunting hut at the head of syttendemajfjorden, ad. s. jensen land, built by nanok in august 1951. fjellborg-hytten 75ø (75°46.0´n 20°08.2´w). norwegian hut built in october 1950 by arktisk næringsdrift on the sw side of langelv, 18 km from the east coast of hochstetter forland. it replaced langelv-hytten. fjerma 73ø (73°10.6´n 23°34.4´w). river on ymer ø draining north into dusén fjord. so named on an nsiu map (1932a). fjordblick schulter 71ø (71°45.0´n 25°48.5´w). snow dome with a view of nordvestfjord in the ne part of the borgbjerg gletscher 169 region, southern stauning alper. probably first climbed and named by the 1977 schwäbische stauning alper expedition. fjord-eidet 71ø (71°38.0´n 22°23.7´w). name used for a norwegian hunting hut said to have been built by the møre expedition in 1931 in nathorst fjord. however, p.s. mikkelsen (1994) records that the first hut built on this site was siste-huset, erected in 1932. (fjordeidet.) fjordbotten – see bundhytten i tyrolerfjord. fjordhytten 74ø (74°38.6´n 20°49.2´w). danish hunting hut on the south side of lindeman fjord, built by nanok in august 1938. it is also known as lindeman fjord hytten. it was burnt down in december 1978 (p.s. mikkelsen 1994). fjortenkilometernæsset – see fyrretyvekilometernæsset. flade isblink 81ø-73 (81°30.0´n 14°30.0´w; maps 1, 4). large, flat ice cap in northern kronprins christian land. mapped and named by lauge koch during reconnaissance flights in 1933 during the 1931–34 treårsekspeditionen (flade = flat). flade teltø 76ø-217 (76°45.2´n 20°59.0´w). island off se daniel bruun land near port arthur. so named by the 1938–39 mørke fjord expedition because of the occurrence of 14 large inuit tent rings (telt = tent). fladebugt 74ø-57 (74°23.7´n 19°09.0´w; maps 2, 4). small bay on the east coast of wollaston forland. named flache bai by karl koldewey’s 1869–70 expedition because the water was so shallow that the greater part of the bay became dry at low tide (flache = flade = flat). the name has appeared on the published geodætisk institut map sheet (74 ø.1) and also on ams map sheets in the form flakkebugt (see also flakkebjerge). (flache bugt, flachen bai, flache bay.) fladebugt 77ø-55 (77°15.0´n 19°15.0´w; map 4). bay on the south side of skærfjorden. so named by david malmquist during the 1931–34 treårsekspeditionen. eigil knuth (1940) expressed the view that this name should never have been given as sledging across the bay was usually difficult due to high snow drifts, and to describe it as flat (= flade) was misleading. fladedal 73ø-431 (73°07.3´n 24°36.1´w; map 4). large, flatbottomed valley in the central part of ymer ø, named by silvio eha during lauge koch’s 1947–49 expeditions. fladedalhytten – see flatdalshytta. fladegletscher 72ø-167 (72°29.2´n 22°04.2´w). glacier on eastern traill ø on the north flank of ellemandsbjerge. so named during lauge koch’s 1936–38 expeditions by hans p. schaub because it is fairly flat. fladepynt 70ø-12 (70°29.7´n 28°35.1´w). low-lying point on the north side of vestfjord. named by carl ryder’s 1891–92 expedition as flade pynt, and traversed on the expedition’s second sledge trip in may 1892. fladestrand – see lapstun hytten. fladstrand 74ø-197 (74°05.7´n 21°13.2´w). beach near eski monæs, south clavering ø, on the east side of østhavn. the name was first used in botanical reports of the 1931–34 treårsekspedi tionen (gelting 1934). flagellarislette 81ø (81°12.5´n 13°32.0´w). plain in central kilen, kronprins christian land. the name records a variety of saxifraga, and is found on a coloured geological map of kilen printed in 1991. flakkebjerge 72ø-172 (72°22.0´n 23°06.4´w). mountain range about 800 m high on southern traill ø, overlooking kong oscar fjord. the name was one of a group given by the place name committee in 1939 (flakke = low, flat). flakkebugt – see fladebugt. flakkerhuk 70ø-255 (70°28.8´n 23°23.2´w; maps 3, 4). flat-lying coastal region in south jameson land, characterised by a moraine ridge system 1–2 km wide and 50–80 m high. hermann aldinger’s original name for this feature was the highway, and it was changed to flakkerhuk by the place name committee in 1935. flata 73ø (73°28.1´n 21°56.8´w). norwegian hunting hut in bad land dal, nw of myggbukta, built by arktisk næringsdrift in 1931 (nsiu 1932c). it was named for the flat terrain. it has also been known as giesecke. flatbreen 73ø (73°33.0´n 29°38.0´w). name used by arne høygaard and martin mehren in 1931 for the present hamberg gletscher. they described it as a large, flat glacier about 10 km wide with a very low gradient. flatdalshytta 73ø (73°02.6´n 24°42.4´w). norwegian hunting hut on the south side of ymer ø at the mouth of fladedal, built in august 1934 by arktisk næringsdrift. it has also been known as firmannsdalen and namsdalsstua. (flatdalen, fladedalhytten.) flatstranda 71ø (c. 71°52´n 24°45´w). norwegian hunting hut built in august 1931 by the møre expedition on the nw side of fleming fjord, and named for the flat terrain around the hut. it was destroyed by high seas in 1953 (p.s. mikkelsen 1994, 2008). (stranda-huset.) flatøyra 73ø (73°02.2´n 22°49.9´w). delta on the north side of geographical society ø. so named on the nsiu maps of lacmann (1937) for its flat nature. fleineset 72ø (72°40.6´n 21°58.1´w). small peninsula in extreme se geographical society ø. so named on the nsiu maps of lacmann (1937) after the locality of the same name in vesterålen, norway. fleming dal 71ø (71°32.1´n 23°01.3´w). name sometimes used by norwegian hunters for the present pingel dal, which drains into fleming fjord (ingstad 1935, 1937). the name has occasionally been used in ornithology reports (e.g. marris & ogilvie 1962). fleming dal hytten 71ø (71°33.1´n 22°58.1´w). norwegian hunting hut in pingel dal (sometimes called fleming dal), built by helge ingstad’s expedition in 1932–33 about 8 km south of the head of fleming fjord. it has also been known as pingel dal hytten and landhuset. fleming fjord 71ø-19 (71°45.0´n 22°48.5´w; maps 3, 4). fjord nw of wegener halvø. named by william scoresby jr. in 1822 as fleming inlet after john fleming [1785–1857], noted for his ‘phi losophy of zoology’ published in 1822. he subsequently became professor at aberdeen and edinburgh. scoresby thought that his fleming inlet might have a connection with his hall inlet (now hall bredning) which would have made jameson land an island. amdrup’s expedition demonstrated that it was a fjord (hartz 1902), although the ‘inlet’ form continued to appear on maps for many years. (flemming inlet, flemming fjord.) fleming fjord hytten 71ø (71°52.2´n 22°45.6´w). norwegian hunting hut on the north side of fleming fjord, 10 km sw of kap biot, built for hermann andresen’s expedition in september 1954. it has also been known by the names lapstun-hytten, fladestrand, surøje and søndre biot. (flemmingfjordhytten, fleming-inlet hytte). fleming fjord nord 71ø (71°52.2´n 22°45.5´w). norwegian hut erected in the summer of 1955 for hermann andresen’s expedition beside lapstun hytten on the north side of fleming fjord, 10 km sw of kap biot. the hut was moved by otto lapstun from nat horst fjord, where it was known as kaares-bu. lapstun had intended to place the hut in ørsted dal, but this was never achieved. flemingfjordhuset 71ø (71°43.2´n 22°43.9´w). small wintering station on the east side of fleming fjord at the mouth of the valley vimmelskaftet. it was built in 1934 during the 1931–34 treårseks pedi tionen, and is also known as vimmelskaftet and kap brown huset. (flemmingfjordhuset, fleming fjord huset.) fleskesvoren 72ø (72°05.0´n 24°55.0´w; map 5). icefall in the upper part of gullygletscher. so named by the 1996 norwegian stauning alper expedition because the symmetrical pattern of the crevasses resembled the crackling on roast pork. flett plateau 73ø-298 (73°37.9´n 23°48.8´w; map 4). plateau about 1500 m high on western gauss halvø. so named during the 1931–34 treårsekspeditionen by gunnar säve-söderbergh after sir john smith flett [1869–1947], a british petrologist who was director of the british geological survey from 1920 to 1935. he 170 had interests in devonian fossils and stratigraphy. (fletts plateau.) flexurdal 71ø-151 (71°58.0´n 23°07.7´w). valley draining north into antarctic havn. named by hans stauber during lauge koch’s 1936–38 expeditions for the curved flexure in the rocks following the valley. flexurebjerg 73ø-294 (73°57.1´n 22°14.4´w). mountain 880 m high in eastern hudson land. named during the 1931–34 treårs ekspeditionen by gunnar säve-söderbergh as flexure hill, because the rocks are folded near the summit. (fleksurfjellet.) fligely fjord 74ø-31 75ø-20a (74°56.0´n 20°37.0´w; maps 2, 4; fig. 15). n–s-trending sound bounding the west side of kuhn ø. named by karl koldewey’s 1869–70 expedition after august von fligely [1810–79], an austrian field marshal and cartographer. fligely was noted for his map making, especially of hungary, and was director of the militärgeographische institut (military geo graphical institute); it was from this institute that julius payer had been granted leave to take part in koldewey’s expedition (j. løve, personal communication 2010). (fliegely fjord, fligelys fjord, fligely-fjord.) fligelyhytten 74ø-f26 (74°59.4´n 20°34.0´w). danish hunting hut on the east coast of fligely fjord, about 8 km sw of kap mosle. it was built by nanok in august 1930. (nordlige fligelyhytten.) flipa 74ø (74°06.5´n 21°17.3´w). small river on south clavering ø draining into vesthavn, equivalent to the present vesterelv. the name is used on an nsiu map (1932a) and the maps of lacmann (1937), and derives from the norwegian dialect word (flipa = whine or whimper). flisane 72ø-n87 (72°39.8´n 22°19.5´w). long narrow island at the east mouth of vega sund. so named on the nsiu maps of lacmann (1937) for the shape. flodskær 73ø-244 (c. 73°08´n 22°48´w). small skerry off eastern ymer ø, named on an nsiu map (1932a) as floskjer (flodskær = flo skjer = tidal skerry). fluoritdal 72ø-144 (72°11.1´n 22°31.3´w). valley in extreme se traill ø on the east side of drømmebugten. named during lauge koch’s 1936–38 expeditions by hans p. schaub for the occurrence of the mineral fluorite. flyveplads – see mestersvig. flyverbjerg 80ø-77 (80°07.8´n 21°49.4´w; map 4; fig. 24). moun tain in south kronprins christian land, south of centrumsø. named during lauge koch’s 1952–53 expeditions by erdhart fränkl, for the pilots of the catalina aircraft which transported the expedition. this was the only mountain which the catalina pilots climbed during the summer (flyver = airman, flyer). flyverfjord 71ø-63 (71°32.5´n 28°00.0´w; maps 3, 4; fig. 41). fjord branching off the south side of nordvestfjord between hinks land and th. sørensen land. mapped by lauge koch during flights in 1932 on the 1931–34 treårsekspeditionen. koch noted it as a tribute to his pilot, n.v. petersen (see also kap jørn). the mouth of the fjord was first seen by alwin pedersen during a long sledge journey from scoresbysund in 1929. flyversø 77ø-142 (77°45.0´n 20°37.0´w; map 4). lake in nord marken. so named by john haller following explorations during lauge koch’s 1955 expedition, because it was possible to land on the lake with a heinkel sea plane in august 1955. the name is a tribute to the pilot (flyver = airman, flyer). flødegletscher 72ø-250 (72°15.3´n 24°29.2´w; map 5). glacier in the north stauning alper, draining east into skeldal. named by erdhart fränkl during lauge koch’s 1950–51 expeditions for the milky colour (fløde = cream). (flöde-gletscher, flode gletscher.) fløelv 73ø-304 (73°26.3´n 21°54.7´w). river draining jakob dal flowing across vestersletten, eastern hudson land. the name was proposed by the place name committee (flø is an old danish expression for a reversed tidal flow in the mouth of a river). fog river 70ø (70°27.5´n 23°02.6´w). minor river in south jameson land flowing into hesteelv. the name was used during the 1931–34 treårsekspeditionen by aldinger (1935), and is assumed to record the common coastal fog in the vicinity while working here. foksa 73ø (73°46.9´n 21°55.7´w). river on the west side of loch fyne, draining the east slope of nordhoek bjerg. so named on an nsiu map (1932a), the name may derive from the norwegian dialect word for drifts of snow. foldaelv 73ø-212 (73°24.5´n 22°03.2´w). river draining the southern giesecke bjerge. named on an nsiu map (1932a) as folda, probably because it flows in the valley the norwegians called folddalen, and reaches the coast close to the norwegian hunting hut known as foldvik. ‘folda’ is a common place name in norway, used for a number of fjords and rivers. folddalen 73ø (73°24.5´n 22°04.5´w). valley in the southern giesecke bjerge. so named on an nsiu map (1932a), possibly after the norwegian hunting hut foldvik built at its mouth. it carries the river known as foldaelv. foldebjerg 72ø-447 (72°29.7´n 27°27.3´w; map 4). mountain about 1400 m high in sw gletscherland. so named by eugène weg mann during the 1931–34 treårsekspeditionen because of the pronounced folds in the rocks. foldegletscher 70ø-83 (70°16.0´n 24°47.5´w). glacier on vol fig. 41. the e–w-trending iceberg-filled flyverfjord between hinks land to the north and th. sørensen land to the south. in the background is nathorst land. the icebergs have been calved from the very productive glacier daugaardjensen gletscher, and have drifted east and southeast before running aground in flyverfjord. the john haller photograph collection, geus archive. flyverfjord nathorst land th. sørensen land hinks land 171 quaart boon kyst draining north to terrassevig. named during the 1931–34 treårsekspeditionen by laurits bruhn, probably for the presence of folded dirt bands in the ice. foldvik 73ø (73°22.6´n 21°41.8´w). norwegian hunting hut on the south side of kap bennet, built by the foldvik expedition in august 1927. this name appears on the nsiu (1932a) map, and commemorates nils foldvik [b. 1892], assistant at the geophysical institute in tromsø, and leader of the 1926–28 expedition. folddalen and foldaelv reach the coast close to the hut. the names bennethytta, giskehytta and giesecké have also been used. foldvik kløft 74ø-94 (74°02.2´n 21°35.2´w). ravine about 6 km east of kap stosch, named by lauge koch’s 1926–27 expeditions as foldvik creek after nils foldvik, leader of the 1926–28 foldvik expedition. see also foldvik. this is an important geological type locality for the foldvik creek formation, but was not shown on any of koch’s maps. according to teichert & kummel (1976), based on information from svend bendix-almgren and tove birkelund, it corresponds to the ravine carrying either river 7 or river 8, most probably river 8. (foldviks kløft.) foldvikfjellet 72ø (72°56.5´n 24°01.3´w). mountain 1120 m high on western geographical society ø. so named on the nsiu maps of lacmann (1937) after nils foldvik (see also foldvik). folly 76ø (76°52.3´n 23°08.7´w). surveying station in eastern dronning louise land used by the 1952–54 british north green land expedition. the name appears on the maps of hamilton et al. (1956). foraarsboplads – see forårsboplads. forbindelsesdal 73ø-445 (73°39.5´n 23°05.0´w). valley on the north side of moskusoksefjord providing a route through to ankerbjergsdal. the name is attributed to heinrich bütler, and arose from his work with lauge koch’s expeditions in the 1950s (forbindelse = connection). forchhammer bjerg 72ø-59 (72°15.5´n 22°52.8´w). mountain about 1350 m high on se traill ø. named by a.g. nathorst’s 1899 expedition as forchhammers berg after johan georg forchhammer [1794–1865], a danish geologist and chemist, and professor of mineralogy and geology at the university of copenhagen from 1831 to 1865. (forchhammer mt., mt. forchhammer, forch ham mer fjellet.) forchhammerdal 72ø-147 (72°13.8´n 22°46.8´w). valley in ex treme se traill ø, east of forchhammer bjerg. the name was given by hans p. schaub during lauge koch’s 1936–38 expeditions. forellenseepingo 72ø (72°32.9´n 23°33.7´w). large pingo beside forelsø, south of karupelv, traill ø. the pingo is 29 m high, 515 m in circumference, and was so named by fritz müller during lauge koch’s 1954–55 expeditions after the adjacent lake forelsø. forelsø 72ø-336 (72°32.7´n 23°39.4´w). large lake south of karupelv, traill ø, in the vicinity of five large pingos. the name was given by fritz müller during his work on the pingos in 1954–55, for the numerous trout (= forel). forhindringsgletscher 73ø-373 (73°49.1´n 25°55.3´w). glacier in north andrée land, partially blocking eremitdal. so named during lauge koch’s 1948–50 expeditions by erdhart fränkl because it hindered progress along the valley. (forhindrings gletscher.) forkastningsdalen 73ø-90 (73°58.7´n 21°21.7´w). valley in home forland. the name appears to have been adopted from the great fault valley of koch (1931), a name used for the valley in which blåelv flows. officially it is said to be a side valley draining into blåelv, but this may be an error. forkastningsdalen 73ø (73°10.0´n 23°20.2´w). valley on eastern ymer ø, west of celsius bjerg, so named on an nsiu map (1932a) because the valley is eroded along a fault, a geologically weak zone (forkastning = fault). forkastningspasset 74ø-366 (74°15.6´n 20°38.1´w). pass between grønnedal and storstrømmen, eastern clavering ø. so named during lauge koch’s 1936–38 expeditions by wolf maync and andreas vischer, because it coincides with a major fault line. forkdal 71ø-404 (71°25.3´n 22°42.8´w). small valley in ne jameson land which forks into two at its head. it drains northwards into passagen. named by katharina perch-nielsen during the 1967–72 ggu scoresby sund expeditions. forkgletscher 72ø-150 (72°16.8´n 22°50.3´w). glacier on se traill ø, at the head of steenstrup dal. so named during lauge koch’s 1936–38 expeditions by hans p. schaub because the glacier divides into two parts. formanden 74ø-315 (74°58.6´n 23°01.5´w). nunatak about 1850 m high sw of the head of grandjean fjord, discovered during the journey by curt teichert and th. johansen along the inland ice margin in 1932. the name records its upstanding character (for manden = the chairman). forposten 71ø-139 (71°01.5´n 21°42.0´w; map 4). cape 680 m high in east liverpool land, named during the 1931–34 treårs ekspeditionen by laurits bruhn (forpost = outpost). forposten 74ø-132 (74°17.5´n 20°39.8´w). mountain 1312 m high on east clavering ø, named by lauge koch’s 1929–30 expeditions in the form avantpost mtns. it was situated in front of the main crystalline mountain range, separated from them by djævlekløften (forpost = outpost). (mt avantpost.) forsblad fjord 72ø-30 (72°25.5´n 25°24.5´w; maps 4, 5). fjord between lyell land and nathorst land, continuing eastwards as segelsällskapet fjord. so called by a.g. nathorst in 1899 after nils jakob forsblad [b. 1874], the master of the expedition ship antarctic. (forsblads fjord.) forsteningskløft 71ø (c. 71°17´n 23°03´w). name used by jensen (1909) in his report on mammals seen during g.c. amdrup’s 1898–1900 expedition, and used for a ravine in jameson land. exact locality uncertain, but possibly in the vicinity of fossilbjerget (forstening = fossil). fortet 70ø-107 (70°42.9´n 22°48.2´w). summit in jameson land, ne of j.p. koch fjeld, so named during the 1931–34 treårseks peditionen by laurits bruhn (fortet = the fort). forårsdal 72ø-143 (72°08.7´n 22°24.9´w). valley in extreme se traill ø, se of drømmebugten. named during lauge koch’s 1936–38 expeditions by hans p. schaub for the attractive springlike setting (forår = spring). (foraarsdal.) forårsboplads 76ø (76°55.8´n 20°18.6´w). descriptive name for the inuit (eskimo) ruins 500 m east of gravelven, corresponding ap proxi mately to the present gravsletten. it was used by the 1906–08 danmark-ekspeditionen in the form foraarsboplads (forår = spring). forårsstedet 73ø-413 (73°59.8´n 28°23.3´w). locality west of eleonore sø, arnold escher land. so named by hans r. katz during lauge koch’s 1951 expedition because plants were found here during his traverse through the nunatak region. the site seems to be a moraine on a glacier at an altitude of 1500 m. (foraars stedet.) fosdalen 73ø-136 (73°53.6´n 20°49.9´w). valley on the north coast of home forland, draining north into gael hamke bugt. the name appears on an nsiu map (1932a) in the form fossdalen, and was given for a waterfall (= foss). river 25 has also been used. fossdal 72ø-519 (72°24.3´n 23°03.1´w). valley on se traill ø draining east into mountnorris fjord. named by geoffrey halliday following botanical work during the 1961 leicester university expedition and 1971 northern universities expedition (foss = waterfall). fossilbjerget 71ø-44 (71°16.2´n 23°02.8´w). mountain 910 m high in jameson land, west of carlsberg fjord. named by g.c. amdrup’s 1898–1900 expedition as fossil-bjerget. the name originated during the exploration of jameson land by otto norden skjöld and henrich deichmann in august 1900, because they found many fossils here. on the 1968 published 1:250 000 scale geo dætisk institut map sheet (71 ø.1) the name is placed in error 172 against a 1010 m high mountain 13 km to the ne. (fossil berg, fossil mountain, mont des fossils.) fossilelv 74ø-102 (74°08.4´n 20°37.5´w). small river on se clavering ø, named by lauge koch’s 1929–30 expeditions in the form fossil river because of the common occurrence of fossils. on norwegian maps it appears as svinta. foster bugt 72ø-15 73ø-272 (73°15.0´n 21°30.0´w; maps 3, 4). broad bay between hold with hope and bontekoe ø. named by douglas clavering in 1823 as foster’s bay, in compliment to henry foster [1796–1831], a midshipman on clavering’s ship who also drew the chart. foster subsequently sailed with william parry on his 3rd and 4th arctic voyages, and was drowned in 1831 during an expedition to panama. (foster bay, fosterbukta, foster-bukta, fosters bugt.) foto sø 74ø-196 (74°06.0´n 21°15.7´w). small lake ne of eski monæs station on southern clavering ø, on detailed maps (1:10 000) a freshwater lake at 42 m altitude draining east into østerelv. a small hut was built here. large numbers of aerial photographs taken by norseman aircraft were developed at eskimonæs during lauge koch’s expeditions. fox havn [ujaattuttalerajiip kangererajiva] 70ø-314 (70°27.9´n 21°56.6´w). harbour south of scoresbysund, south liverpool land. named after the fox ii, a 409 ton steam-driven barque that, renamed grønland, carried the scoresbysund colonisation expedition in 1924 and lost its rudder when almost wrecked near fox pynt. the ship was later renamed gustav holm and sailed regularly to east greenland, notably with lauge koch’s geological expeditions. it was sold in 1951 and broken up. the fox ii was originally built for the kryolite company in 1893, as a replacement for the more famous fox used by f.l. mcclintock in 1857–59 on his search voyage for sir john franklin’s lost expedition. fox lake 76ø (76°15.1´n 18°41.5´w). lake on store koldewey where sampling was undertaken for phytoplankton studies (cremer et al. 2005). fox pynt [qattiterpaajik] 70ø-313 (70°28.2´n 21°56.7´w). penin sula between amdrup havn and fox havn, on the east side of rosenvinge bugt, south liverpool land. named after the ship grønland, formerly the fox ii, which was almost wrecked near this point in 1924. see also fox havn. foxtrap point 74ø (74°05.5´n 21°17.1´w). name used in the archaeology report of j.m. wordie’s 1926 expedition for the peninsula eskimonæs, because two well-preserved inuit fox traps were found here. framnes 74ø (74°05.6´n 21°05.8´w). peninsula east of eskimovig in south clavering ø. named after the norwegian engineer and hunter niels framnes hansen, who hunted in east greenland from 1928 to 1930 and 1935 to 1937, and in southern east greenland from 1931 to 1933. the name was first used on nsiu maps (lac mann 1937). frankfurter spids 71ø (71°49.6´n 25°23.2´w). mountain on the se side of the upper basin of spærregletscher, eastern nathorst land. climbed by karl m. herligkoffer’s 1966 expedition on 19 august, and named after the german city of frankfurt. the map in fantin (1969) is difficult to fit with modern detailed maps. franklin dal 73ø-441 (73°16.4´n 22°10.7´w). steep valley near kap franklin, se gauss halvø. named during lauge koch’s 1950 expedition by p. graeter. see also kap franklin. franklin strand hytten – see franklinhytta. franklin ø 72ø-16 (c. 72°39´n 21°39´w). small island off geo graphical society ø. the name franklin island first appeared on the 1872 admiralty chart no. 2282, and according to white (1927) was probably a mistake by the draughtsman, who may have had kap franklin in mind when engraving the copper plate. franklin-huset 73ø (73°20.4´n 21°57.5´w). norwegian hunting station built for johan a. olsen’s expedition between kap franklin and kap bennet. it was demolished by the foldvik expedition in 1927 and the material used to build franklinhytta and foldvik. franklindalen 73ø (73°20.3´n 22°14.5´w). name occasionally used in norwegian reports (e.g. bang 1944) for a valley in the southern giesecke bjerge, the present randbøldalen. björnedalen has also been used. franklinfjellet – see franklinspitze. franklinhytta 73ø (73°18.3´n 22°05.6´w). norwegian hunting hut about 7 km north of kap franklin, southern giesecke bjerge, built by the foldvik expedition in september 1927 using material from the 1922 franklin-huset nearby. it appears as franklinstranda on an nsiu map (1932a). (franklin-hytta, franklin-stranda, franklin strand hytten.) franklinspitze 73ø (73°17.1´n 22°18.4´w). mountain about 1200 m high behind kap franklin, southern giesecke bjerge, now known as knuden. the name was used in koldewey’s (1874) narrative of his 1869–70 expedition, in the description of the first ascent by ralph copeland and julius payer on 8 august 1870. an nsiu map (1932a) uses franklinfjellet. see also kap franklin. (franklinspitze, franklinfjæld.) franske øer 78ø-4 (78°40.0´n 18°20.6´w; maps 1, 4). island group east of jøkelbugten, north of the pariserøerne. the duke of orléans in 1905 had given the name îles françaises to an island group south of the pariserøerne, approximating to the position of the present danske øer (fig. 9). the 1906–08 danmark-ekspedi tionen transferred the name to the present position, corresponding to the northernmost islands the orléans expedition could have seen. the wac maps from 1952 retain the name in the original position. (franske islands.) frebold bjerg 73ø-43 74ø-205a (74°00.6´n 21°37.0´w; map 4). mountain 1207 m high in home forland, sw of kap stosch. named by lauge koch’s 1929–30 expeditions as mt frebold, after hans frebold, a geologist who studied permian and carboniferous fossils from the 1930 collections and took part in the 1931–34 treårsekspeditionen. frederiksborg gletscher – see øvre frederiksborg gletscher. frederiksborg nunatakker 69ø-47 (69°02.0´n 31°45.0´w). group of nunataks between lindberg fjelde and prinsen af wales bjerge, named by l.r. wager’s 1935–36 expedition as frederiksborg nuna taks after frederiksborg castle in denmark. frederiksdal 71ø-148 72ø-436 (71°53´n 26°40´w to 71°38´n 26°29´w; map 4). valley system extending from the west end of furesø south to trianglen, then east to reach nordvestfjord at nordbugt. named during the 1931–34 treårsekspeditionen by ove simonsen after the danish locality of the same name nw of copenhagen (see also furesø). fredhaug 75ø (75°57.9´n 20°48.2´w). norwegian hunting hut built for john giæver’s expedition in september 1932 on the south side of bessel fjord. it has also been known as svarthammerhytten. freeden bugt 74ø-368 75ø-29a (75°00.0´n 18°00.0´w; map 4). bay on the south side of shannon. named freeden bai by karl kolde wey’s 1869–70 expedition after wilhelm ihno adolph von freeden [1822–94], a german teacher of navigation and founder and director of norddeutschen seewarte (north german naval obser vatory). he was one of the principal supporters of koldewey’s expedition, and wrote the introduction to the meteorology and hydrography chapters of the expedition narrative (verein für die deutsche nordpolarfahrt in bremen 1873–74). (freeden bay, freedenbukta, freedens bucht, freedensbucht.) freja fjeld 74ø-296 (74°49.8´n 21°10.8´w). mountain about 1300 m high in th. thomsen land. the name originated from the wintering parties at eskimonæs and kulhus during the 1931–34 tre årsekspeditionen, and was given for freja, daughter of njord, goddess of love in nordic mythology and the greatest of the female gods. frejagletscher 74ø-379 (74°24.0´n 20°52.5´w). glacier on north clavering ø draining into skilledal. named in the form fröjabreen 173 on the nsiu maps of lacmann (1937) after the goddess freja. see freja fjeld. the name was approved in 1950, after it had been used in reports of glaciological studies. (frøya glacier, fröya glacier.) freke 74ø-294 (74°45.0´n 21°17.5´w). mountain 844 m high at the south end of odin dal, th. thomsen land. the name originated from the wintering party at kulhus during the 1931–34 treårs ekspeditionen, and was given for one of odin’s dogs in nordic mythology. freuchen gletscher 71ø-375 (71°27.8´n 29°38.0´w; map 4). gla cier between royston nunatakker and sw hinks land, draining se into krummedal. named by peter vogt during lauge koch’s 1957 expedition, after peter freuchen [1886–1957], danish journalist, author and arctic explorer associated with knud ras mussen’s thule expeditions. freuchens hytte 76ø (76°55.3´n 21°01.6´w). name commonly used for the meteorological station in pustervig manned by peter freuchen during the 1906–08 danmark-ekspeditionen. see also pustervig. freyberg fjæld 70ø (70°35.7´n 22°40.2´w). mountain 673 m high on the west side of hurry inlet, south of astartekløft. the name was used in a report by hermann aldinger (1935) on work during the 1931–34 treårsekspeditionen, and was given for bruno von frey berg, a german geologist who had worked in brazil. (freyberg fjæld.) freycinet bjerg 72ø-10 (72°46.7´n 22°24.3´w; map 4; fig. 12). mountain about 900 m high on eastern geo graph ical society ø. william scoresby jr. named cape freycinet in 1822 after louisclaude desaules de freycinet [1779–1842], a french navigator who made notable voyages of discovery to australia between 1800 and 1804, and sailed around the world in the l’uranie in 1817–20. scoresby’s cape was identified by white (1927) as a mountain on geographical society ø and renamed mount freycinet, later frey cinet bjerg. (freycinetfjellet.) fribjergene 74ø-335 (74°01.7´n 24°14.7´w; map 4). mountain 1300 m high in ole rømer land, so named during lauge koch’s 1936–38 expeditions by heinrich bütler. it was named after the 17th century swiss national park in glarus, friberg, where the animals range at liberty (= fri). frieda sø 76ø (76°06.5´n 20°14.7´w). lake in ad. s. jensen land at the head of syttendemajfjord, the present gunnar andersen sø. the name was given for j.g. jennov’s wife and was reported as in common use among danish hunters. it first appears on the maps of the 1932 gefion expedition, and appeared periodically in later publications (e.g. jennov 1963). repeated attempts by jennov to obtain official approval of the name were unsuccessful. (friedasø.) friedas ø 75ø (76°06.7´n 20°53.3´w). name reported used by danish hunters from about 1923 for the present ulla ø at the mouth of grandjean fjord. it was named after j.g. jennov’s wife. see also frieda sø. friggbreen 74ø (74°19.5´n 21°06.0´w). glacier on central clavering ø, draining east into skillegletscher. so named on the nsiu maps of lacmann (1937) after frigg, wife of odin in old nordic mythology. frihedsgletscher 72ø-315 (72°11.7´n 25°03.6´w; map 5). tribu tary glacier on the north side of vikingebræ, west of friheds tinde, stauning alper. named by john haller following explorations during lauge koch’s 1954 expedition (frihed = freedom, liberty). see also frihedstinde. frihedspas 72ø-497 (72°12.2´n 24°59.8´w; map 5). pass about 1800 m high between vikingebræ and skjoldungebræ, north of friheds tinde, stauning alper. named by the 1963 cambridge university expedition. frihedstinde 72ø-251 (72°11.7´n 24°58.1´w; map 5; figs 27, 42). mountain 2610 m high in the north stauning alper, between the heads of vikingebræ and skjoldungebræ. the first ascent was made by peter braun and fritz schwarzenbach in august 1951, and the name was given to honour the freedom-fighters of the danish resistance during world war ii (frihed = freedom). this ascent has been claimed as the first major climb in the stauning alper. fritz johansen ø 76ø-347 (76°19.0´n 21°14.5´w; map 4). island north of ad. s. jensen land. named by john haller following ex plorations during lauge koch’s 1956–58 expeditions, after fritz johansen [1882–1957], zoologist on the 1906–08 danmark-eks pedi tionen. he subsequently took part in vilhjalmur stefans son’s 1913–18 arctic expedition. frique peak 72ø (72°05.3´n 24°37.2´w; map 5) mountain a short distance ne of glamis borg in the northern stauning alper. the name was used by the 1991 scottish stauning alper expedition, which failed to reach the summit. (frique.) froggies beaut 72ø (72°04.2´n 24°52.2´w; map 5). peak on the divide between the heads of gullygletscher and schuchert glet scher. the name is used by the 1996 norwegian stauning alper expedition in their report and on their maps, but was not claimed as a first ascent. frosnebugt 75ø-28 (75°07.1´n 17°44.7´w; map 4). large bay on east shannon. named by karl koldewey’s 1869–70 expedition as gefrorne bai, because the bay was still ice-covered and impassable when their ship reached here in july 1870 (gefrorne = frosne = frozen). (gefrorene bai, frozen bay.) frydendal 71ø-136 (71°01.8´n 22°07.5´w). valley on the south side frihedstinde fig. 42. looking west to frihedstinde, a 2610 m high peak in the stauning alper. the john haller photograph collection, geus archive. 174 of storefjord, central liverpool land. so named by laurits bruhn during the 1931–34 treårsekspeditionen (fryd = delight). frænkel land 73ø-514 (73°18.0´n 27°35.0´w; maps 3, 4). land area bounded by isfjord and inner kejser franz joseph fjord. originally named as frænkels halfö by a.g. nathorst’s 1899 expedition, after knut hjalmar ferdinand frænkel [1870–1897], a swedish engineer and member of salomon andrée’s balloon expedition to the north pole on which he made the meteorological observations. nathorst’s expedition was searching for traces of the lost andrée balloon expedition. (frænkel peninsula, fränkels land.) fuchs bjerg 73ø-62 (73°42.6´n 22°37.9´w). mountain ridge up to 1600 m high in eastern hudson land, named by lauge koch’s 1929–30 expeditions in the form fuchs ridges or mt fuchs after vivian ernest fuchs [1908–1999]. he was one of the geologists of j.m. wordie’s 1929 expedition that made investigations in this region. fuchs was director of the falkland islands dependencies survey from 1947 to 1950 and 1960 to 1973, and is particularly noted for his leadership of the 1955–58 transantarctic expedition. (fuchsfjellet.) fuchsberg 71ø (c. 71°06´n 24°12´w). locality in west jameson land, about 5 km ne of alfred wegener’s 1930–31 eastern scientific station, where a fox den was observed on the summit of a small hill (fuchs = fox). (fox hill.) fugleneset 74ø (c. 74°16´n 19°23´w). name used by severin liavaag’s 1908–09 expedition (brandal 1930) for a feature in the vicinity of kap borlase warren, se wollaston forland, where the expedition had shot ptarmigan (fugl = bird). exact position uncertain. fuglenæbsfjeldet 76ø-51 (76°57.8´n 20°33.1´w). mountain 810 m high in daniel bruun land, north of mørkefjordsbugten. so named by the 1906–08 danmark-ekspeditionen for a resemblance to a bird’s beak. koch (1916 p. 398) used the latin form rostrum avis on the christmas card sent to peter freuchen at pustervig in 1907. (fuglenæbsfjeld, fuglenæb mt.) fuglesø 76ø (76°46.4´n 18°43.4´w). name reported by fischer (1983) as used by staff at danmarkshavn for lille skibssø, a small lake sw of vandsø. a hide apparently built by alwin pedersen for bird-watching in 1938–39 was said to be still in use in the early 1980s. fuglesøen 73ø-662 (73°18.7´n 25°06.1´w). small lake on the south side of noa dal, ymer ø. the name originated during the 1931–34 treårsekspeditionen, presumably for the numerous birds seen there, and was approved at the suggestion of r. spärck. fuglevarden 72ø-204 (72°13.8´n 23°46.2´w; map 5). highest point of the peninsula hovedet, east of the mouth of noret, north scores by land, marked by a mound built up from generations of bird droppings (fuglevarden = the bird cairn). named by prospecting teams associated with lauge koch’s 1948–49 expeditions. fugleø 70ø-393 (70°25.0´n 27°48.2´w). small island on the north side of inner føhnfjord. named by the 1963 geodætisk institut expedition for the many birds. fugleø 76ø-273 (76°51.7´n 20°22.4´w). small island in nw dove bugt, east of vædderen. named by the 1938–39 mørkefjord expedition for the numerous birds. fugleøya 72ø (72°47.9´n 22°54.1´w). island in vega sund, the present gåseøen. this name appears on the nsiu maps of lac mann (1937), and derives from the abundant traces of geese and ducks seen here. (fugleöya.) fulach gletscher 72ø-452 (72°53.5´n 26°56.7´w; map 4). glacier on the north side of dickson fjord, suess land. so named by eugène wegmann during the 1931–34 treårsekspeditionen, after fulach near schaffhausen, switzerland. he visited the glacier in 1932. fulachtal 72ø (72°53.5´n 26°56.7´w). name used by andersen (1937), and probably intended for the valley in southern suess land containing fulach gletscher. on his map it is misplaced southwards to gletscherland. fundal, funddal – see nedre funddal and nordre funddal. fundal glacier 72ø (72°06.7´n 24°05.4´w). term used by pessl (1962) for the glacier formerly occupying the valley he calls fundal (nedre funddal), near mestersvig. funkhütte 75ø (75°19.2´n 17°48.1´w). in 1943 a german meteorological station was established at kap sussi on the east coast of shannon. the expedition lived at different times on their ship the coburg, in a camp on the ice, in a cavern excavated in a snow fan, and in a hut on land known as funkhütte. the station was destroyed by american forces in the summer of 1944, but many relics of the station remain, as well as the grave of gerhard zacher, shot by the sledge patrol on 22 april 1944. funkis – see firkanten and syveren. fura 73ø (73°29.0´n 21°21.3´w). river on the south coast of hold with hope, east of myggbukta. named in this form on an nsiu map (1932a; fig. 13), possibly after the river of the same name in the hedmark district of norway. furesø 72ø-95 (72°00.8´n 26°00.0´w; maps 3, 4). e–w-trending 30 km long ice-dammed lake in nathorst land. named during the 1931–34 treårsekspeditionen by ove simonsen after the danish lake of the same name nw of copenhagen (see also frederiksdal). it is dammed at the east end by spærregletscher, where hans gsellman reported a rise in water level of 1.3 m in 48 hours in 1957 when the outlet was blocked by ice. furkla 74ø (74°07.8´n 20°49.4´w). stream in a steep ravine on the east side of dødemandsbugten, south clavering ø. used on an nsiu map (1932a) and maps of lacmann (1937), the name is derived from the norwegian dialect word for a ravine. furnes 74ø (c. 74°42´n 20°08´w). norwegian hunting hut on southern kuhn ø, 3 km west of kap hamburg, erected by the møre expedition in august 1930. it was named after jørgen furnes [b. 1897], a norwegian hunter who overwintered in east greenland from 1927 to 1929. the hut was moved to this site from kap schumacher where it was known as agnes-tufta, and at the present location has also been known as kap hamburghytten, røsnes and kapp norge. now disappeared. furnesfjellet 74ø (74°07.4´n 21°00.0´w). mountain about 900 m high on south clavering ø, equivalent to the present jernhatten. so named on the nsiu maps of lacmann (1937) after jørgen furnes – see also furnes. füssener ryggen 71ø (71°48.3´n 25°02.1´w; map 5). ridge on the sw side of roslin gletscher, stauning alper. it was climbed by karl m. herligkoffer’s expedition on 20 august 1966, and named after füssen, a small picturesque town in the bavarian alps, germany. fynselv 70ø-102 (70°31.9´n 23°14.4´w; map 4). river in jameson land flowing south into scoresby sund. so named by laurits bruhn during the 1931–34 treårsekspeditionen after the island of fyn, denmark. fyrbøderdal 69ø-28 (69°45.0´n 23°22.7´w). this is probably a valley on turner ø at the east side of turner sund, although the precise location is uncertain. the name was used in the form fyrböderdal in böggild’s (1905) report on mineral collections from g.c. amdrup’s 1898–1900 expedition (fyrbøder = stoker). fyriselv 77ø-66 (77°31.3´n 20°38.2´w). river in east nordmarken draining south into h.g. backlund fjord. mapped in 1933 by david malmquist during the 1931–34 treårsekspeditionen, and named after the swedish river fyrisån that he passed every day on the way to his office when drawing up the map. fyrretyvekilometernæsset 77ø-17 (77°01.7´n 18°11.0´w; map 4). peninsula on the east coast of germania land. so named by the 1906–08 danmark-ekspeditionen, because it was approximately 40 km sledging distance from the expedition base at danmark havn. the published diaries of poulsen (1991) and thostrup (2007) demonstrate that numerous other capes on the east coast of ger 175 mania land were given informal names recording the approximate sledging distance from danmark havn; these include niog halv tredskilometernæsset (59 km), fireogtyvekilometernæsset (24 km), fjor ten kilometernæsset (14 km), niogtredivekilometernæsset (39 km), seksogtredivekilometernæsset (36 km), and treogtredive kilometer næs set (33 km). none of these capes can be precisely located on mo d ern maps. nioghalvtredskilometernæsset was also known as snefogsdepot according to poulsen (1991). fyrtårnet 72ø-208 (72°10.8´n 23°56.1´w; map 5). one of the summits of korsbjerg nw of mesters vig. named by prospecting teams associated with lauge koch’s 1948–49 expeditions (fyrtårnet = the beacon). (fyrtaarnet.) fældestrand 80ø-8 (80°23.5´n 15°45.8´w; map 4). stretch of coast on the east coast of holm land, south of eskimonæs. so named during the 1906–08 danmark-ekspeditionen by christian b. tho s trup because of the presence of large, very well-preserved inuit fox traps. (shore of traps.) fønbugt 70ø-406 (70°31.0´n 26°56.5´w). bay on the north side of fønfjord. named by w. stuart watt during the 1967–72 ggu scoresby sund expeditions after fønfjord (føn = katabatic wind). føndal 72ø-135 (72°08.1´n 22°15.4´w). small valley in the extreme se of traill ø. the name was adopted from a suggestion by arne noe-nygaard during the 1931–34 treårsekspeditionen, and de rives from the strong winds experienced here in 1932. fønfjord [ujuaakajiip kangertiva] 70ø-18 (70°28.0´n 27°00.0´w; maps 3, 4). fjord between milne land and gåseland. named by carl ryder’s 1891–92 expedition as føhnfjord because strong winds were encountered here on their first day of exploration in august 1891 (fig. 7). the variation blastfjord was used by gulløv (1991); blast = blæst = wind (j. løve, personal communication 2010). (föhnfjord, føhn fiord, føhn fjord.) första nålbrevet 71ø (71°04.0´n 25°31.9´w). name used in a report by helge g. backlund on work during the 1931–34 treårseks peditionen for the north pinnacle on bjørneøer vi, a surveying point on one of the islands climbed in 1933. første hvide 74ø-169 (74°21.1´n 20°32.0´w). one of three lightcoloured areas of sedimentary rocks in ne clavering ø, contrasting with dark basalts. they were named by arne noe-nyegaard and gunnar säve-söderbergh during the 1931–34 treårseks peditionen, and first appeared in the forms erste weisse, zweite weisse and dritte weisse, normally used in danicised form as 1. hvide, 2. hvide and 3. hvide. see also anden hvide and tredie hvide. første hytten 73ø (73°38.9´n 23°10.5´w). norwegian hunting hut built for arktisk næringsdrift in august 1932 on the north side of moskusoksefjord. it was subsequently renamed petrahytten, and has also been known as røiskattlia. førstemai-bukta 73ø (73°24.0´n 25°16.3´w). small bay near kap petersen in nw ymer ø, where two norwegian hunters, john giæver and søren richter, camped on 1 may 1930 during a journey from blomsterbugt to eleonore bugt. g g. glacier – see gerard de gerr gletscher. gabet [nuukajiit akornganni kangerterajik] 70ø-218 (70°40.4´n 21°38.8´w). bay on the east coast of liverpool land, between hagen and snuden. named during the 1931–34 treårseks pe di tionen by laurits bruhn for its shape on the map (gabet = the jaws). (gabet bugt.) gadekæret 74ø (74°28.2´n 20°34.0´w). locality in the vicinity of zackenberg forskningsstation. the name is used as a reference locality by visiting scientists. gael hamke bugt 73ø-1, 74ø-90a (74°05.0´n 19°53.0´w; maps 2, 4; fig. 15). large bay between clavering ø and home forland. the present position corresponds with that of the baey door gael hamkes found on the 1666 dutch charts of hendrick doncker and peter goos, and said to have been discovered by a dutch skipper of that name in 1654. scoresby (1823) had positioned the bay incorrectly, and the present position is that determined by clavering (1830). (gael hamkes bay, gael-hamkes-bucht, bay of gael hamkes, gael hamke bai, gaél hamke fjorden, t'bay v. gale hamkes, baay van gale heinkes). gaffeldal 73ø-50b (73°58.2´n 21°22.7´w). minor valley on the north slope of stensiö plateau, draining from the east into blåelv, nw hold with hope. so named by eigil nielsen during the 1931–34 treårsekspeditionen because the valley splits into many branches (gaffel = fork). gaffelelv 70ø-169 (70°41.3´n 22°25.5´w). river in south liverpool land with two main tributaries, draining west into hurry inlet. named during the 1931–34 treårsekspeditionen by laurits bruhn (gaffel = fork). gaffelfjeld 70ø-168 (70°42.6´n 22°15.2´w). mountain ridge about 600 m high in southern liverpool land, south of sødal, drained by gaffelelv. so named by laurits bruhn during the 1931–34 treårs eks peditionen. gaffelgletscher 72ø-155 (72°18.4´n 22°31.7´w). glacier on se traill ø, south of mountnorris fjord. named during lauge koch’s 1936–38 expeditions by hans peter schaub for its fork-like shape (gaffel = fork). gaflen 80ø (80°33.3´n 19°40.4´w). glacier on the west side of the prinsesse caroline-mathilde alper, inner ingolf fjord, which forks upwards into two branches. named by the 1938–39 danske hunde slæde-ekspedition (drastrup 1945) for its fork-like shape (gaffel = fork). the name is also found on 1957 ams maps. galadriel fjeld 81ø (81°13.6´n 13°55.9´w). hill 356 m high in central kilen, kronprins christian land. the name is found on a coloured geological map of kilen printed in 1991 (pedersen 1991), and is said to derive from tolkien’s ‘lord of the rings’. gale hamke’s land 74ø (c. 74°30´n 19°30´w). name appearing on a number of old dutch maps, e.g. on hendrick doncker’s 1663 chart as landt door gaal hamkes and joh. van keulen’s 1681 map as landt van gaal hamkes, and said to have been discovered by a dutch skipper of that name in 1654. see also gael hamkes bugt. it was the land most frequently reported as having been seen by whalers in the 17th and 18th century, who probably saw part of the present wollaston forland or the pendulum øer. william scoresby had placed the name at about latitude 75°n in 1822. use of the name was discontinued in the 1930s by a decision of the place name committee. galenadal 72ø-241 (72°17.8´n 25°29.3´w; maps 4, 5). valley in east nathorst land on the west side of alpefjord, named by erdhardt fränkl during lauge koch’s 1950–51 expeditions for finds of galena-bearing quartz veins. veins in this area also contain other ore minerals (harpøth et al. 1986). gamle jim øer 79ø-41 (79°21.3´n 19°22.1´w; maps 1, 4). island group on the east side of lambert land, one of five names given by the place name committee after dogs used on the 1906–08 danmark-ekspeditionen. the dog ‘gamle jim’ appeared to have died during a three-day snow storm on a sledge journey, but revived when kicked (gamle = old). gamle jonsbu – see jónsbú. gamma havn 76ø-291 (76°55.6´n 20°18.0´w). small bay in front of mørkefjord station where the gamma anchored to unload equipment for the 1938–39 mørkefjord expedition. the gamma was a three-mast, 200 ton danish schooner built in 1919 at thorseng, and purchased and strengthened for the expedition. (gamma harbour.) gamma ø 77ø-94 (77°50.0´n 19°49.0´w; maps 1, 2, 4). large island north of orléans sund. named by the 1938–39 mørkefjord expedition after the expedition ship gamma. see also gamma havn. orleans island has also been used. 176 gammel axels tinde – see akselborg. gammel hellerup gletscher 78ø-36 (78°33.0´n 21°41.4´w; map 4). glacier between nørreland and nørre mellemland, hertugen af orléans land. named during the 1938–39 mørkefjord expedition by svend sølver, after gammel hellerup gymnasium, his old school. gammen – see borganes and germania-hamn. gamssteig 75ø (c. 75°19´n 17°48´w). feature in the vicinity of the base camp of the 1943–44 operation bassgeiger at kap sussi, shannon. the name is recorded by olsen (1965). gamvik 73ø (73°38.7´n 20°28.3´w). norwegian hunting hut on the east coast of hold with hope, nw of holland ø. built by nils foldvik’s expedition in august 1927, the hut name appears in the orvin (1930) list of hunting huts. it may have been named after gamvik in the tromsø region of norway. gannochy gletscher 71ø (71°48.0´n 24°37.5´w; map 5). glacier in the stauning alper, between storgletscher and roslin gletscher. named and explored by the 1968 university of dundee expedition. garagebugt 71ø (71°35.0´n 27°58.0´w). bay on the north side of flyverfjord, the present lancaster bugt, where numerous icebergs are stranded, as if parked in a garage. so named during the 1931–34 treårsekspeditionen by helge g. backlund (in: koch 1955). garbh bheinn 72ø (72°07.4´n 24°31.7´w). name appearing in an early report of malcolm slesser’s 1958 expedition (bennet 1959) for the present dunvegan toppene, in the north stauning alper. it was named for its resemblance to a scottish mountain of the same name. garde nunatakker 78ø-19 (78°28.0´n 22°29.0´w; maps 1, 4). nuna tak group west of hertugen af orlèans land, including grønne nunatak and tuborgfondet land. named by the 1909–12 alabama expedition after captain thomas vilhelm garde [1859–1926], director of the naval department and from 1918 admiral, who had showed great kindness and interest in the expedition. it was garde who had granted iver iversen leave to join ejnar mikkelsen’s expedition (j. løve, personal communication 2009). garde was particularly noted for his participation in the 1883–1885 umiak expe di tion to southern east greenland with gustav holm. (garde’s nuna takker.) garmischer spids 71ø (71°47.5´n 24°59.1´w; map 5). mountain 2209 m high on the sw side of roslin gletscher, stauning alper. climbed by karl herligkoffer’s 1966 expedition, and named after the bavarian mountain resort of garmish-partenkirchen. gasserfjellet 72ø (72°55.7´n 23°13.0´w). mountain about 1000 m high on central geographical society ø. used only on nsiu maps (lacmann 1937), the name was given for max gasser [1872–1954], one of the pioneers in the development of practical mapping techniques from aerial photographs. gastisdal 73ø-70 (73°30.0´n 22°39.6´w). valley on gauss halvø draining north into moskusoksefjord. named by lauge koch’s 1929–30 expeditions in the form gastis valley, it commemorates the then famous uppsala restaurant ‘gästis’. gauche peak 72ø (72°02.0´n 24°50.1´w; map 5). mountain at the head of schuchert gletscher, se of trumpington pas, stauning alper. first climbed by the 1961 university of bangor expedition. gauli dal 73ø-695 (73°20.0´n 26°13.1´w). valley in southern andrée land, draining west from gauli gletscher. named by john haller following explorations during lauge koch’s 1949–51 expeditions. see also gauli gletscher. gauli gletscher 73ø-695a (73°20.3´n 26°09.1´w). glacier in southern andrée land. named by john haller following explorations during lauge koch’s 1949–51 expeditions after the glacier of this name in the bernese oberland, switzerland, where an american war plane had made an emergency landing. gauli gletscher in andrée land seemed also to be a possible emergency landing site. gausa 73ø (73°37.9´n 21°45.5´w). river in west hold with hope flowing into the head of loch fyne. so named on an nsiu map (1932a), possibly after the river gausa in the oppland district of norway. gauss halvø 73ø-31 (73°30.0´n 23°00.0´w; maps 3, 4). peninsula between muskusoksefjord and kejser franz joseph fjord. karl koldewey’s 1869–70 expedition originally gave the name cap gauss to a point on the south side of this peninsula, probably the present sydvestpynt, but a.g. nathorst’s 1899 expedition was unable to determine the position because of the rounding of the coast and applied the name gauss halfö to the entire peninsula. it was named after karl friedrich gauss [1777–1855], a german mathematician, astronomer and physicist. (gauss peninsula, gauss halvöya, gauss halbinsel, gaußhalbinsel.) gavlen 70ø-267 (70°06.3´n 23°30.2´w). mountain 1150 m high on volquaart boon kyst, so named during the 1931–34 treårseks peditionen by laurits bruhn because in shape it resembled the gable of a house. gedderyggen 74ø-285 (74°15.0´n 21°51.1´w). mountain ridge 1050 m high on west clavering ø. the name originated from the wintering party at eskimonæs during the 1931–34 treårseks peditionen and was given for the spikey appearance, like the fin of a fish (gedde = pike). gefion havn 76ø-160 (76°23.0´n 20°53.6´w). harbour on the south coast of godfred hansen ø, sw dove bugt. so named after the three-masted danish schooner gefion, the ship of the 1932 gefion expedition which anchored here and unloaded building material in the harbour. the nanok hunting station ålborghus was built here in 1938. (gefions havn, gefionhavn.) gefion havn hytten 76ø (c. 76°23´n 20°54´w). danish hunting hut built for nanok in may 1934 at gefion havn, on the south side of godfred hansen ø, sw dove bugt. it was replaced in 1938 by ålborghus hunting station. gefion pass – see øvre gefionpas. gefiontinder 76ø-133 (76°28.1´n 25°39.0´w). group of summits in sw dronning louise land, named by j.p. koch’s 1912–13 expedition as gefions tinder. gefion was the virgin sister of danish gods said to have ploughed out the island of sjælland from sweden in a single night. (gefionstinder, gefionland, gefions-zinnen, gefju nar tindar.) geheimrat finsterwalders fjell 72ø (72°57.4´n 23°33.7´w). n–strending mountain ridge in geographical society ø. used on the nsiu maps of lacmann (1937), the name was given for sebastian finsterwalder [1862–1951], a german pioneer of theoretical de vel opments in aerial photogrammetry. geikie plateau 69ø-30 (69°55.0´n 26°00.0´w; maps 3, 4). extensive ice plateau south of scoresby sund. it was mapped and named by lauge koch during flights in 1933 on the 1931–34 treårseks pe di tionen. the name commemorates archibald geikie [1835–1924], a scottish igneous petrologist and stratigrapher, professor of geology at the university of edinburgh from 1871 to 1881, and director of the british geological survey from 1882 to 1901. geisha 73ø-259 (73°27.2´n 21°01.3´w). norwegian hunting hut on the south coast of hold with hope, 15 km east of myggbukta, built by the foldvik expedition in august 1926. it was named after their alsatian dog ‘geisha’ purchased in tromsø for the expedition. (geischa.) gelbe rinne 71ø (71°57.5´n 24°16.7´w). zone of prominent red and yellow staining surrounding the molybdenum deposit at malm bjerg, eastern stauning alper (gelb = yellow, rinne = furrow). the name arose during prospecting by nordisk mineselskab (harpøth et al. 1986). gemini col 71ø (71°50.4´n 25°30.6´w). col between a ne branch of borgbjerg gletscher and the southern part of spærregletscher. this high pass was traversed by the 1988 scottish staunings expedition. gemmedal 73ø-645 (73°39.7´n 27°04.6´w; map 4). valley in west andrée land draining sw to gerard de geer gletscher. so named 177 during the 1931–34 treårsekspeditionen by ove simonsen be cause it is well hidden and difficult of access (gemmedal = hidden valley). general director river 70ø (70°30.0´n 22°53.8´w). name used by hermann aldinger during the 1931–34 treårsekspeditionen for the present mønselv, a river in southern jameson land. genvejsdalen 73ø-327 (73°45.6´n 23°34.8´w). valley in moskus okselandet, south hudson land, draining into moskusoksefjord. the name is a modification of a suggestion by heinrich bütler arising from his work during lauge koch’s 1936–38 expeditions. the valley provided a route to the interior of hudson land (genvej = short cut). see also hurtigrute-tal. geographical society ø 72ø-64 73ø-277 (72°57.0´n 23°30.0´w; maps 3, 4). large island between sofia sund and vega sund. named by a.g. nathorst’s 1899 expedition as geographical society’s ö after the royal geographical society of london, because of its great interest in arctic research. the society had also made a contribution to nathorst’s expedition. (geographical society’s island, geographical societyöya, geographical society-øya, geographic society insel.) geologfjord 73ø-517 (73°45.0´n 25°18.0´w; maps 2–4). fjord be tween strindberg land and andrée land. named geolog fjorden by a.g. nathorst’s 1899 expedition because of the spectacular and colourful rock formations, and in honour of his own pro fession. (geologists fjord, geology fjord.) geologhytten 73ø (c. 73°34´n 24°52´w). norwegian hunting hut on the west side of geologfjord, east of mørkebjerg, andrée land, built by arktisk næringsdrift in september 1933. disappeared. it was also known as mørkebjerghytten and brandalhytten. geologists glacier 74ø (74°42.8´n 22°45.2´w). tributary glacier to pasterze, th. thomsen land. the name was used informally by battle (1952), a tribute to the first man known to have sledged down it in 1938–39, the swiss geologist adolf ernst mittelholzer. gerard de geer gletscher 73ø-570 (73°34.0´n 27°15.0´w; maps 2, 4). large n–s-trending glacier between louise boyd land and andrée land. it was named by louise boyd and carl-julius anrick in 1931, originally as g. glacier, subsequently de geer glacier (anrick 1932), after the swedish geologist gerard baron de geer [1858–1943]. de geer was noted for his six expeditions to spits bergen, where his explorations gave rise to many of the place names. he was the originator of the varve-counting method of glacial geochronology, and professor of geology at the university of stockholm from 1897 to 1924. (g. glacier.) gerda gletscher 74ø (74°41.0´n 22°36.3´w; fig. 1932geodætisk insititut). name used on the 1932 edition of the geodætisk institut 1:1 million scale map for the present pasterze. on this map pasterze was moved to the position of the present copeland gletscher. gerda gletscher was said to have been named by lauge koch after the danish actress gerda madsen. gere 74ø-293 (74°45.7´n 21°21.1´w). mountain 902 m high at the south end of odin dal, th. thomsen land. the name originated from the wintering party at kulhus during the 1931–34 treårseks peditionen, and was given for one of odin’s dogs in nordic mythology. germania bjerg 74ø-2 (74°33.3´n 18°47.6´w). mountain 302 m high on southern sabine ø. named by karl koldewey’s 1869–70 expedition as germaniaberge after the expedition ship germania (fig. 43), a 90-foot, 143-ton steamer built at bremerhaven in 1869 for the expedition. the mountain may correspond to scoresby’s cape bright. (mt. germania, germaniabjerget.) germania ekspeditionens varde 77ø-19 (c. 77°04´n 18°56´w). cairn in central germania land erected by karl koldewey’s expedition on 15 april 1870, and marking their farthest north. it was found by the 1906–08 danmark-ekspeditionen who described it as two feet high and built on an inconspicuous summit. koldewey’s message was illegible. exact position uncertain. germania havn 74ø-47 (74°32.2´n 18°49.9´w; map 4). small en closed bay on the south side of sabine ø. this was the winter harbour of the germania (fig. 43), karl koldewey’s 1869–70 expedition ship. see also germania bjerg. the original name for the bay was germaniahafen. edward sabine conducted pendulum experiments on the shore of the bay in 1823, and koldewey’s expedition carried out astronomical observations in 1869–70. the danish hunting station germaniahavn was built here in 1919. (ger mania hafen, germaniahamnen, germania har bour.) germania land 76ø-11 77ø-110a (77°00.0´n 19°00.0´w; maps 2, 4). large land area between skærfjorden and dove bugt, so named by the 1906–08 danmark-ekspeditionen. j.p. koch (1916) records that the name was given by mylius-erichsen to commemorate its fig. 43. the germania was the ship that carried karl koldewey’s 1869–70 second german north pole expedition to northern east greenland. from: verein für die deutsche nordpolarfahrt in bremen (1873–74). 178 discovery by karl koldewey’s 1869–70 expedition in the ger mania (fig. 43), and was also intended as a compliment to alfred wegener, the german member of the 1906–08 danmark-eks peditionen. see also germaniahavn. current approved usage re stricts the name to the area east of valdemarsmuren. (germania land, germania halbinsel). germania land hytten 77ø (77°01.0´n 19°05.8´w). hut built by danmarkshavn weather station personnel in 1979, wnw of ger mania ekspeditionens varde in germania land. germania-hamn 74ø (74°32.1´n 18°51.0´w). norwegian hunting hut built in 1909 at germania havn, southern sabine ø, by vebjørn landmark. it was also known as gammen. a danish hunting sta tion was built nearby in 1919. see also germaniahavn. germaniahavn 74ø (74°32.2´n 18°48.3´w). danish hunting station built in 1919 at germania havn in southern sabine ø by øst grønlandske fangstkompagni. the station was manned in the periods 1919–20, 1921–24 and 1928–31; it was also known as ger maniahavn-huset, blæsebælgen and villaen (p.s. mikkelsen 2008). in 1948 it was replaced by a new station built by nanok, which is still maintained by sirius. ruins of earlier huts include a norwegian station built in 1909, and karl koldewey’s observatory dating from 1869. a hut beside the station known as h.l. jensens hus was taken down in 1923. (germania havn station.) germaniahavn-huset – see germaniahavn. gessnerfjellet 74ø (74°13.6´n 20°58.9´w). mountain on southern clavering ø. so named on the nsiu maps of lacmann (1937) after wilhelm gessner [b. 1890], director of hansa luftbild gesell schaft, berlin, which undertook construction of the detailed nsiu maps of east greenland. ghiacciaio brescia 70ø (70°05.3´n 23°02.4´w). minor glacier west of milano gletscher on the northern blosseville kyst. named by leonardo bonzi’s 1934 expedition after the north italian city of the same name. (brescia glacier.) ghiacciaio genova 70ø (70°01.4´n 23°15.9´w). glacier above klin ten on volquaart boon kyst, draining south, corresponding in part to the present torvgletscher. the glacier was first traversed during leonardo bonzi’s 1934 expedition, and was named after the italian city of the same name. on some italian maps (fantin 1969), torv gletscher is placed south of latitude 69°n and extends to the coast, and ghiacciaio genova is shown as a tributary to it on its northern side. (genova glacier.) gibson’s point 70ø (70°35.5´n 22°26.0´w). prominence in hurry inlet named by william scoresby jr. in 1822 as gibson’s point or point gibson, after one of his two partners on the baffin. the name is not marked on his chart, although it can be identified (scoresby 1823, p. 463) as a point on the west coast of liverpool land, the present suluppik south of the mouth of gubbedal. giesecké 73ø (73°22.6n 21°41.8´w). norwegian hunting hut south of kap bennet, built by the foldvik expedition in 1927. this name appears on an nsiu map published in 1929, and was given for its situation east of the giesecke bjerge. the hut has also been known as foldvik, bennethytta and giskehytta. giesecke – see flata. giesecke bjerg 74ø-66 (74°28.8´n 21°46.7´w; map 4). mountain 1328 m high on the north side of tyrolerfjord, southern a.p. olsen land. named by karl koldewey’s 1869–70 expedition as cap giesecke, after the german naturalist karl ludwig (charles lewis) giesecke [1761–1833], who made extensive mineral collections in west greenland between 1806 and 1813, and from 1813 was professor of geology at the university of dublin, ireland. it is the mountain that has the appearance of a cape, but this is not so clear on a map, and the name was therefore applied by the place name committee to the mountain forming the ‘cape’. (gieseckes bjerg.) giesecke bjerge 73ø-8 (73°27.0´n 22°07.0´w; map 4). range of mountains in eastern gauss halvø. william scoresby jr. on his 1822 voyage had named cape giesecké in compliment to charles lewis giesecke [1761–1833] – see also giesecke bjerg. a few early danish maps placed kap giesecke south of mackenzie bugt at the present site of kap bennet. nathorst (1901) suggested the name be given to a mountain, which was probably what scoresby had seen. j.m. wordie’s 1926 expedition extended the name to the range of mountains between kap franklin and ladder bjerg. (cape gie secke, giesecké mountains, giesecke bjærge, giezecke-fjella, giske landet.) gieseckedalen 74ø (74°28.2´n 21°39.3´w). name used by norwegian hunters for the valley on the north side of tyrolerfjord east of giesecke bjerg, a.p. olsen land. a hut at the mouth of the valley used for fishing was known as giskehuset. (giesecke dal.) gieseckehytten – see giskehuset. gilbert-murray brae 72ø (72°06.1´n 26°53.5´w). small glacier near the mouth of jomfrudal, nathorst land. the name was introduced by geoffrey halliday during the 1961 leicester university expedition, and was given for gilbert murray, a pioneer of british rock climbing. one of the halls of residence at leicester university is called after him. gille valley 73ø (73°30.5´n 22°52.2´w). valley on gauss halvø, draining north to moskusoksefjord. so named by gunnar sävesöderbergh during the 1931–34 treårsekspeditionen, after the then famous restaurant in uppsala, sweden. gimle 71ø-185 (71°31.2´n 23°46.5´w). mountain 928 m high in northern jameson land, nw of olympen. the name was one of a group of names given by the place name committee in 1939 to replace proposals by hans stauber. ‘gimle’ in norse mythology was the golden hall where (after ragnarok) the good would enjoy eternal happiness. gimli height 73ø (73°32.6´n 25°45.9´w). summit 2062 m high on the south side of grejsdalen, andrée land; described as a fine summit with a knife-edge ridge. climbed by the 2007 army boreal zenith expedition. gipsdalen 71ø-162 (71°49.5´n 23°43.2´w; map 4). valley south of the werner bjerge draining south and east into ørsted dal. named during lauge koch’s 1936–38 expeditions by hans stauber for the gypsum-bearing triassic rocks. girton fjeld 72ø-513 (72°01.7´n 25°00.7´w; map 5) mountain between gully gletscher and sefström gletscher, stauning alper. bennet (1972) placed the mountain immediately se of churchill pass. named by the 1963 cambridge university expedition, which made the first ascent on 21 august, after girton college, cam bridge. a noted women’s college originally founded at hitchin in 1869, girton college was transferred to cambridge in 1873. (girton.) giskehuset 74ø (74°27.1´n 21°41.9´w). norwegian hunting hut se of giesecke bjerg, a.p. olsen land, built by the w. holmboe fishing expedition in 1932 for salmon fishing. it is also known as holmboehytten and bjørnestua. (gieseckehytten.) giskehytta – see giesecké. giskeodde 73ø (73°23.4´n 21°35.5´w). name sometimes used by norwegian hunters for kap bennet, eastern gauss halvø, derived from its position east of the giesecke bjerge. gisvold 74ø (74°25.6´n 20°20.9´w). norwegian hunting hut in sw wollaston forland, on the ne side of zackenberg bugt. it was built by nils foldvik’s expedition in 1927, and named after arnulf gis vold, a member of the expedition. it has also been called norske pashytten. givskovselv 74ø (74°09.5´n 20°36.3´w). river on east clavering ø, the present moskusokseelv. the name appears on a sketch map in gustav thostrup’s 1921 logbook (møller 1939), and was apparently given for the danish hunter, hans givskov. giæver-hytta 73ø (73°42.2´n 24°30.6´w). norwegian hunting hut in strindberg land at the mouth of brogetdal, erected by arktisk næringsdrift in 1930. it was transported from a site on the south side of gauss halvø on two sledges. named after john schelderup 179 giæver [1901–70], who used this hut from 1930 to 1931, and led his own expedition to regions farther north in 1932–34. he was a journalist from 1921 to 1929, and then became an arctic trapper visiting east greenland, the white sea and jan mayen. giæver was secretary of nsiu in 1935, and from 1937 to 1956 (apart from the war years and 1949–52) leader of the annual summer relief expeditions. the hut has also been called strindberghytta. it was demolished and the material used to build strindberghuset. (giæver-tun, giæverhytten.) giæverdalen 73ø (73°45.8´n 24°48.8´w w). norwegian hunters name for brogetdal, the large valley in strindberg land which drains eastwards into nordfjord. the name appears on several nsiu maps and in hunting accounts from about 1932, and is still occasionally seen in norwegian publications. strindberg valley has also been used. see also giæver-hytte. glacier 21 72ø (72°05.8´n 24°28.4´w). temporary name used for the present kishmul gletscher in the stauning alper in early reports of malcolm slesser’s 1958 climbing expedition (bennet 1959). glacier bj. petersen 70ø (70°35.0´n 21°51.2´w). minor glacier north of scoresbysund, southern liverpool land, se of trefoden. the name was used on an inaccurate 1933 chart by m. parat prepared during j.b. charcot’s expeditions, and named after bjerring pedersen. the name was not approved, because another glacier in southern liverpool land had already received the name bjerring pedersen gletscher. charcot’s 1925 expedition on the pourquoi pas? had sent the first report of pedersen’s death during the 1924–25 colonisation expedition back to denmark. glacier ch. maurain 70ø (70°46.0´n 25°57.5´w). small glacier on east milne land, a minor tributary to charcot gletscher on its north side. the name was used by parat & drach (1934) in their report on j.b. charcot’s 1933 expedition, and was named after chevallier maurain, a french professor who had participated in the 1932 expedition. glacier chatton 70ø (70°45.0´n 25°46.5´w). glacier on east milne land corresponding to the present charcot gletscher. the name was used by parat & drach (1934) in their report on j.b. charcot’s 1933 expedition. see also chattonbugt. glacier de la petite sirène 71ø (71°55.3´n 25°48.0´w). minor glacier on the east side of prinsessegletscher, eastern nathorst land, named and traversed by claude rey’s 1968 expedition during their ascent of pic ludovica. (sirène = siren). glacier dérobé 71ø (71°53.1´n 25°50.6´w). minor glacier on the west side of prinsessegletscher, eastern nathorst land. named by claude rey’s 1968 expedition. glacier des lutins 71ø (71°57´n 25°48´w). minor glacier on the east side of prinsessegletscher, west stauning alper. named by claude rey’s 1968 expedition (lutin = troll). glacier des myrtilles 71ø (71°58.2´n 25°55.2´w). tributary glacier on the west side of prinsessegletscher, eastern nathorst land. named by claude rey’s 1968 expedition (myrtille = bilberry). glacier des oubliettes 71ø (71°55.8´n 25°56.0´w). tributary glacier on the west side of prinsessegletscher, east nathorst land. named by claude rey’s 1968 expedition, perhaps for the crevasses and cavities within the ice (oubliette = dungeon). glacier des sires d’equealoir 71ø (71°57.0´n 25°54.0´w). minor glacier on the west side of prinsessegletscher, east nathorst land. named by claude rey’s 1968 expedition. glacier des tours 71ø (71°58.5´n 25°47.7´w). glacier east of prin ses se gletscher, western stauning alper, named by claude rey’s 1968 expedition which traversed the glacier during their climb of tour vercors and tour chartreuse (tour = tower). glacier des violettes 71ø (71°52.4´n 25°50.3´w). tributary glacier on the sw side of prinsessegletscher, eastern nathorst land. named by claude rey’s 1968 expedition, probably for the colour of the ice (violette = purple, violet). glacier du furesoe 71ø (71°57.0´n 25°50.5´w). name used by claude rey’s 1968 expedition for the present prinsessegletscher, which drains north into furesø. glacier du renard 71ø (71°52.0´n 25°42.1´w; map 5). minor tributary glacier on the east side of prinsessegletscher, western stauning alper. so named by claude rey’s 1968 expedition, presumably for the sighting of a fox (= renard). glacier j.l. faure 70ø (70°40.9´n 26°04.0´w). glacier tongue draining south into vinkeldal, se milne land. the name was used by parat & drach (1934), and named after jean-louis faure, a french surgeon who accompanied j.b. charcot’s 1932 expedition and wrote an account of that voyage (faure 1933). glacier lauge koch 70ø (70°34.0´n 21°47.8´w). glacier in south liverpool land ne of scoresbysund, draining north to lillefjord. the name was used on maps and in accounts of the ‘campagne du pourquoi pas?’ led by j.b. charcot (e.g. faure 1933; parat & drach 1934). the french expeditions had received help and advice from lauge koch, and considered him one of denmark’s most eminent geologists. see also lauge koch bjerg. (glacier lauge kock.) glacier le mouchoir 71ø (71°54.5´n 25°45.5´w). minor glacier on the east side of prinsessegletscher, western stauning alper, named and traversed by claude rey’s 1968 expedition on their ascent of pic ludovica. the name may recall its small size (mouchoir = handkerchief ). glacier watkins 70ø (70°37.2´n 21°51.2´w). name used on an inaccurate 1933 map drawn by m. parat during j.b. charcot’s 1933 expedition for a small glacier on the west side of lillefjord, southern liverpool land. see also baie watkins. glamis borg 72ø-368 (72°05.0´n 24°39.2´w; map 5). mountain 2200 m high between bersærkerbræ and kishmul gletscher, northern stauning alper. first climbed by malcolm slesser’s 1958 expedition, and named after glamis castle, angus, the imposing and historic 17th century home of the earls of strathmore and king horne, and said to be queen elizabeth ii’s favourite castle. the second ascent was by guido monzino’s 1963 expedition that called it cima di granita, and the third ascent by toni gobbi’s party in 1967. (glamis.) glamis gletscher 72ø-370 (72°04.6´n 24°41.5´w; map 5). minor glacier on the se side of bersærkerbræ, sw of glamis borg, north stauning alper. named glamis glacier by malcolm slesser’s 1958 expedition. glamis pas 72ø-369 (72°04.7´n 24°38.6´w; map 5). col between glamis gletscher and the head of kishmul gletscher, north stauning alper. the approved position of the pass is se of glamis borg, the name having originated from malcolm slesser’s 1958 expedition who climbed the mountain from the pass. in most mountaineering literature (e.g. bennet 1972), glamis col (72°05.8 24°34.9w; map 5) is placed on a lower col ne of glamis borg. glasgow ø [tartaajik] 70ø-235 (70°48.6´n 21°39.1´w). small island off the coast of liverpool land, named by william scoresby jr. in 1822 as glasgow island after the scottish city. (île glasgow, glasgow ö.) glatze 71ø (71°55.0´n 25°41.5´w; map 5). snow mountain on the east side of prinsessegletscher, western stauning alper, at the head of castor glacier and pollux glacier. named and first climbed by the 1967 berchtesgaden expedition. glaukonitbjerg 70ø-46 (70°40.0´n 25°17.1´w). minor summit about 180 m high nw of kap leslie, east milne land. named by hermann aldinger during the 1931–34 treårsekspeditionen as glaukonitberg or glaukonit berg, for the presence of the mineral glauconite in the sandstones. gleditschfjellet 72ø (72°55.5´n 23°20.8´w). mountain about 1200 m high on geographical society ø. the name was used only on nsiu maps (lacmann 1937), and commemorates kristen gran gle ditsch [1867–1946], a norwegian colonel who was head of norges geografiske opmåling (norwegian geographical survey). glemmedalen 72ø-304 (72°02.2´n 23°47.9´w; map 5). valley on 180 the ne side of werner bjerge, draining into the head of oksedal. so named by peter bearth and eduard wenk during lauge koch’s 1953–54 expeditions because the valley is hidden from sight until one is abreast of its mouth (glemme = forget). glesdalen 74ø (74°18.4´n 19°49.0´w). valley in southern wollaston forland, west of herschell bjerg, corresponding to the present blæsedalen. the name appears on an nsiu map (1932a), and may derive from a glistening appearance. gletscherbugt 70ø-214 (70°39.2´n 21°46.2´w). bay or small fjord on the se coast of liverpool land, a nw branch of lille fjord. so named by laurits bruhn during the 1931–34 treårsekspeditionen for the glacier draining into the bay. it has also been called baie watkins. gletscherdal 73ø-296 (73°27.5´n 23°05.9´w). small valley on gauss halvø draining nw to join paralleldal. named by lauge koch’s 1929–30 expeditions. the original usage was dobbeltglacier valley (seidenfaden 1931), a mixture of danish and english, and a reference to the glaciers occupying the valley. gletscherland 72ø-426 (72°40.0´n 27°00.0´w; maps 3, 4). land area bounded by dickson fjord and wahlenberg gletscher, and divided almost into two parts by röhss fjord. the name was adapted from a suggestion by ove simonsen during the 1931–34 treårs ekspeditionen, and was given for the many ice caps and glaciers. cantonsland has also been used. gletscherpas 73ø-353 (73°56.9´n 25°11.3´w). pass in central strindberg land at the south end of alpedal, in front of a glacier that drains both nw and se. named during lauge koch’s 1948–49 expeditions by hans r. katz. gletscherpingo 72ø (72°33.5´n 23°31.8´w). name used by fritz müller during lauge koch’s 1954–55 expeditions for a pingo beside karupelv, traill ø. the pingo is 350 m across and 38 m high, and the ice-core bears a close resemblance to glacier ice (müller 1959). gletscherryggen 72ø-183 (72°08.8´n 24°16.1´w; map 5). ridge in north scoresby land, east of skeldal. named by prospecting teams associated with lauge koch’s 1948–49 expeditions for the glacier on the flank of the highest, southernmost summit. gletscherskærene 79ø-19 (79°44.2´n 17°439.9´w; map 4). small skerries off the coast of east hovgaard ø, south of a glacier lobe draining the ice cap which just reaches the coast. named by the 1938–39 mørkefjord expedition. gletschersø 74ø-392 74°16.7´n 25°02.3´w; map 4). lake at the front of the south branch of korsgletscher, southern bartholin land. named by john haller following explorations during lauge koch’s 1956–58 expeditions. glipa 73ø (73°27.5´n 22°04.6´w). river in the giesecke bjerge draining the present gustav dal. the name appears only on the nsiu (1932a) map, and may derive from the norwegian word for a long, small opening. glòës sø 76ø-229 (c. 76°58´n 21°36´w). easternmost lake in vigfus dal, daniel bruun land, west of the head of mørkefjord. so named by the 1938–39 mørkefjord expedition after j.p. koch’s best dog (glòë) which accompanied him on his crossing of the inland ice in 1912–13. exact location uncertain, as the ‘lake’ appears to be one of the wide stretches of the river. glommen 73ø-151 (73°33.1´n 20°49.8´w). river in se hold with hope, named on an nsiu map (1932a; fig. 13) in the form glåma. a common place name in norway, it is probably derived from the dialect word meaning milky water. glückstadt nunatak 77ø-50 (77°09.3´n 24°58.8´w; fig. 21). nuna tak in nw dronning louise land, named by the 1909–12 ala bama expedition after consul general valdemar josef glückstadt [1868–1942], a member of the alabama expedition committee. (glückstadt’s nunatak). glyphea elv 70ø (70°29.2´n 22°12.5´w). name used by alfred rosen krantz (1942) for a river draining south from gulfjelde in south liverpool land. it was named after the fossil glyphea. (glypheaelv.) glöysa 73ø (73°23.8´n 23°11.4´w). stream on the south side of gauss halvø, flowing in the present aina dal. so named on an nsiu map (1932a). (gløysaa.) gneisdal 73ø-674 (73°38.0´n 26°24.0´w; map 4). western branch of grejsdal, central andrée land. so named by john haller following explorations during lauge koch’s 1949–51 expeditions, be cause of the westward change from metasediments to high grade paragneisses along the valley. gnejsnæs 79ø-36 (79°01.9´n 20°49.5´w; maps 1, 4). peninsula in sw lambert land protruding into zachariae isstrøm. named by john haller following explorations during lauge koch’s 1956–58 expeditions, for the rock type (gnejs = gneiss). gnejssø 70ø-381 (70°15.0´n 29°13.0´w; map 4). lake in western gåseland, at the west end of vindblæsedal. so named during lauge koch’s 1958 expedition by eduard wenk because of the gneissic rocks around the lake. lauge koch had landed on the lake during a reconnaissance flight in august 1957. (gneiss sø.) gnipa-höhlen-gletscher 76ø (76°47.8´n 18°45.8´w). small glacier nw of danmark havn in which the ice cave gnipahulen is excavated. the name is used in the koch & wegener (1911) scientific report of the 1906–08 danmark-ekspeditionen. see also gnipa hulen. (gnipa cave glacier, höhlengletscher.) gnipahulen 76ø-233 (76°47.8´n 18°45.8´w). extensive ice cave nw of danmark havn, which periodically collapses and re-forms. it was so named by the 1906–08 danmark-ekspeditionen and described and illustrated by koch & wegener (1911). trolle (1909) described it as a castle of ice, a cathedral of colour and light. named after gnipahelleren, a cave in norse mythology. jennov (1935) reported that the cave had completely melted away before his visit in 1932, while thomsen (1966) visited it in 1950 and penetrated the cave system for 250 m, and fischer et al. (2009) report a visit in 1980. it is reported to have collapsed again in 1988. (gnipahöhle, gnipa-höle, gnipa grotto, gnipa cave.) gnisten 73ø (73°26.9´n 20°38.1´w). norwegian hunting hut built in september 1947 in se hold with hope at kap broer ruys. stein sørensen, who erected the hut, was telegraphist (= gnisten = the spark) at the myggbukta hunting and weather station. gnitaheia 74ø (74°17.7´n 20°52.8´w). mountain ridge in central clav ering ø on the east side of skillegletscher. so named on the nsiu maps of lacmann (1937), after a character in the german epic poem from c. 1200, the nibelungenlied. godfred hansen ø 76ø-167 (76°27.0´n 20°54.5´w; map 4). island in sw dove bugt, where the ålborghus hunting station was established in 1938. so named during the 1932 gefion expedition after godfred hansen [1867–1937], an officer in the danish navy who took part in amundsen’s gjoa expedition from 1903 to 1906 and the 3rd thule expedition 1919–20. he was chairman of øst grønlanske fangstkompagni nanok. the name was not approved until after his death in 1937. godthåb golf 74ø-146 (74°08.0´n 21°53.0´w; map 4; fig. 15). inner embayment south of clavering ø, divided from gael hamke bugt by the finsch øer. named by lauge koch’s 1929–30 expeditions in the form godthaab golf, evidently after the godthaab, which served as expedition ship on koch’s 1929–1933 and 1937 voyages. the godthaab was a 287-ton barquentine built at sande fjord in 1898, and purchased on the slip by grønlands styrelse. she made more than 60 voyages, mainly to east greenland as an expedition ship and as a supply ship to ammassalik and scoresbysund, and was laid up in 1951. clavering fjord, clavering sund and inner bay have been used for the same stretch of water. (godthaab rhede, godthaabs golf, godthåbs golf ). gog 73ø-534 (73°14.4´n 28°24.7´w; map 4; fig. 65). mountain about 2600 m high in west frænkel land, one of two similar high mountains west of the head of knækdal known as gog and magog. they were named by j.m. wordie’s 1929 expedition, probably after 181 the gogmagog hills near cambridge. gog and magog were two giants whose wooden effigies guard the guildhall in london, and were the supposed survivors of a race of legendary giants. the first ascent of gog was made by n.e. odell and his wife during louise boyd’s 1933 expedition, the second by john haller’s party in 1951, and the third by a ggu party in 1975. gog magog glacier 73ø (73°15.0´n 28°19.7´w). name used by odell (1937a, b), for the glacier between the mountains gog and magog, west frænkel land. gondulfjellet 73ø (73°06.8´n 23°42.5´w). mountain ridge about 1530 m high on ymer ø, south of dusén fjord. named in this form on an nsiu map (1932a), possibly for its shape (gondol = gondola). goniomyakløft 70ø-142 (70°35.2´n 22°36.1´w). ravine in neill klinter on the west side of hurry inlet. named by alfred rosen krantz during lauge koch’s 1926–27 expeditions as goniomya kløft after the numerous fossils of the lamellibranch goniomya. (goniomyaklöft.) gonville fjeld 72ø-506 (72°05.1´n 25°11.1´w; map 5). caius fjeld and gonville fjeld are two rock summits about 2280 m high on the west side of cavendish gletscher, stauning alper. they were first climbed by the 1963 cambridge university expedition, and named after gonville and caius college, cambridge. the college was founded by edmond gonville in 1348 and re-founded by dr. caius in 1557. goodenough land 72ø-411 (72°55.0´n 28°20.0´w; maps 3, 4). land area between nordenskiöld gletscher and kjerulf fjord. the name first appears on the 1932 1:1 million scale geodætisk institut map prepared on the basis of 1932 aerial observations by lauge koch during the 1931–34 treårsekspeditionen. it was given for admiral goodenough, then president of the royal geographical society of london. the area was partially explored and mapped by j.m. wordie in 1926 and 1929. a number of the place names here were given after geologists from geneva, switzerland (fritz schwar zen bach, personal communication 1996). (goodenoughs land.) goose cliff 74ø (74°09.7´n 20°11.7´w). reference locality used by madsen (1925) for a breeding site of barnacle geese. rosenberg et al. (1970) suggest it was located at kap mary, eastern clavering ø, or perhaps at basaltkap, southern clavering ø. goose lake 76ø (76°26.6´n 18°48.2´w). lake on store koldewey where sampling was undertaken for phytoplankton studies (cre mer et al. 2005). gorm spids 72ø-195 (72°12.1´n 24°04.6´w). mountain on the east side of store blydal, north scoresby land. named by prospecting teams associated with lauge koch’s 1948–49 expedition, after gorm den gamle, danish king of part of jylland from 936–940, with his seat at jelling. (gorms spids.) government station 72ø (72°13.9´n 23°55.1´w). designation occasionally used in reports and maps for the airfield now known as mestersvig (e.g. washburn 1965). graah bugt / graah gletscher – see below. double ‘a’ (aa) is treated as å in danish. graben land 71ø-438 (71°09.0´n 28°50.4´w; maps 3, 4). large nunatak area between eielson gletscher and vindue gletscher. named by peter homewood during the 1967–72 ggu scoresby sund expeditions for the faults which characterise the area and form a geological structure known as a graben. granatbjerg 73ø-700 (73°14.6´n 27°06.7´w; map 4). mountain about 2100 m high in frænkel land. named during lauge koch’s 1949–51 expeditions by john haller, for the abundant garnets, which reach up to 40 cm in diameter near the summit. granatdal 74ø-262 (74°12.4´n 21°24.5´w). valley on sw clavering ø in which granatelv flows. the name came into use about 1935, and was given for the presence of garnets in the rocks. greindalen has also been used. granatelv 74ø-172 (74°12.4´n 21°24.5´w). river on sw clavering ø, flowing in granatdal. the name was first used in reports of the 1931–34 treårsekspeditionen (malmquist 1932), and records finds of garnets. (garnet fluss, granatelva.) granathytta 74ø (74°09.3´n 21°31.4´w). norwegian hunting hut on the coast of south clavering ø, west of the mouth of granatdal. the hut was built by the foldvik expedition 5 km farther east in 1926, and moved to the present site in 1927. it has also been known as sandvik, svampebugthytten, and as granittelva, a corruption of granatelv. now a ruin. granatskæret 76ø-276 (76°35.3´n 20°43.2´w). small island north of andreas lundager ø, western dove bugt. it was named by the 1938–39 mørkefjord expedition, presumably for the presence of garnets in the rocks. grande jorasses 72ø (72°05.2´n 24°51.8´w). mountain 2750 m high at the head of bersærkerbræ, north stauning alper, equivalent to the present c.f. knox tinde. so named by malcolm slesser’s 1958 expedition because of a resemblance to the mountain of the same name sw of chamonix. it was first climbed by the 1963 cam bridge university expedition led by c.f. knox, and later the same year by the imperial college expedition. grande jorasses is the name used in most mountaineering accounts. grandjean fjord 74ø-182 75ø-45a (75°00.0´n 21°28.8´w; maps 2, 4; see also fig. 86). fjord between c.h. ostenfeld land and th. thom sen land. mapped and named by lauge koch during flights in 1932 during the 1931–34 treårsekspeditionen, and partially ex plored by gunnar seidenfaden in august 1932. it was named after commander, later captain, emil valdemar asger grandjean [1889–1948], chief of the danish naval air force from 1925 to 1941. (grandjeans fjord.) grandjeanhytten 75ø-106 (75°01.6´n 21°28.1´w). danish hunting hut on the north side of central grandjean fjord about 7 km ne of mågenæs, built by nanok in 1934. only the foundations now remain (1988). (grandjean bundhytte.) granit spids 71ø (71°48.6´n 24°59.0´w; map 5). mountain about 2159 m high on the sw side of roslin gletscher, stauning alper. climbed by karl herligkoffer’s 1966 expedition on 20–21 august, and named for the granitic rocks of which the mountain is formed. it has also been called hird star. granite mountain 73ø (73°47.3´n 22°06.5´w). name used occasionally by koch (1930), apparently for the present nordhoek bjerg (or possibly part of the nørlund alper) in hudson land. named for the occurrence of granite. granite valley 73ø (73°47.9´n 22°43.2´w). name used by helge g. backlund (1930) for the present stordal, because of the considerable developments of granites at its lower end. the name was subsequently used in a more restricted sense by seidenfaden (1931) and backlund (1932) for the side valley to stordal now known as dybendal. the name for the river in this side valley (granitelv) was approved. (granit-tal, granittal, granit valley.) granitelv 73ø-59 (73°47.9´n 22°43.2´w). river in hudson land flowing into stordal. the name was originally used during lauge koch’s 1929–30 expeditions for the lower part of the river now known as storelv, as well as the tributary in dybendal to which it is now applied. see also granite valley. (granit river, granit fluss, granittelva.) granitelv 74ø-403 73ø-351 (73°59.6´n 25°50.0´w). river draining granitsø, nw strindberg land, and flowing into the head of geologfjord. named during lauge koch’s 1948–49 expeditions by hans r. katz for the widespread developments of granite. granitfjeld 74ø-393 (74°18.5´n 24°57.3´w). mountain 1800 m high east of korsgletscher in bartholin land, formed of granite. named by john haller following explorations during lauge koch’s 1956–58 expeditions. granitsattelbeerg – see sadelberg. granitsø 74ø-401 (74°02.9´n 25°41.7´w; map 4). lake in nw strindberg land, named by hans r. katz during lauge koch’s 182 1948–49 expeditions for the outcrops of granite. (granitsee.) grannitelva – see granathytta. granta–kirk passet 72ø (72°00.0´n 25°00.1´w; map 5). pass be tween kirkbrae and grantabrae, about 2100 m high. so named by the 1996 norwegian stauning alper expedition, which had been aiming for the grantalang col but took a slightly too northerly route. grantabotn 74ø (74°18.0´n 22°20.4´w). norwegian hunting hut on the north side of inner grantafjord, southern payer land, built by arktisk næringsdrift in september 1931 (nsiu 1932c). (granta fjord hytten, grantahytta, granta-botn.) grantabrae 71ø (71°59.1´n 25°03.2´w; map 5). tributary glacier on the north side of upper sefström gletscher, so named by the 1963 cambridge university expedition. see also grantafjord. grantafjellene 74ø (74°19.8´n 22°20.1´w). name occasionally used by norwegian hunters for the mountains on the north side of gran ta fjord in southern payer land. grantafjord 74ø-85 (74°18.1´n 22°14.9´w). fjord west of clav ering ø, the name was given by j.m. wordie for the river gran ta (also known as the cam) which runs through the city of cam bridge, england. (grant’s fjord.) grantafjordhytten – see grantabotn. grantagletscher 74ø-163 (74°20.0´n 22°55.0´w; map 4). branch of wordie gletscher draining into grantafjord. named by lauge koch’s 1929–30 expeditions in the form granta gletscher. grantalang col 71ø (71°59.3´n 24°59.3´w; map 5). col on the south side of upper langgletscher (now storgletscher) leading to the head of grantabrae and sefström gletscher, stauning alper. the name was used by bennet (1972). grantapynt 74ø-86 (74°18.1´n 22°03.0´w). elongate peninsula between grantafjord and copeland fjord, west of clavering ø. one of the fixed points in j.m. wordie’s 1926 survey of the region was located here, which he named granta point. graue schlucht – see gråkløft. graupa 74ø (74°07.0´n 21°19.1´w). stream on south clavering ø draining into the sea west of falskenæs. used on the nsiu maps of lacmann (1937), it apparently derives from the norwegian dialect word meaning to dig or excavate. gravelven 76ø-246 (76°56.1´n 20°19.6´n). small river in sw germania land between rypefjeldet and brystet. so named by the 1906–08 danmark-ekspeditionen (thostrup 1911), because an inuit grave was found on one of the small islands in the river delta. it was occasionally called ruinelven (j. løve, personal communication 2009). (grave river.) graven 74ø-333 (74°07.9´n 24°18.0´w). deep valley in ole rømer land draining into the head of krumme langsø. named by hein rich bütler during the 1936–38 two-year expedition (graven = the grave). gravhøjen 72ø-142 (72°55.0´n 23°34.3´w). mountain about 1500 m high ne of lumskebugten, se suess land. named by ove simon sen during the 1931–34 treårsekspeditionen for its appearance (gravhøj = burial mound). mount marcel bertrand has also been used. gravsletten 76ø-295 (76°55.8´n 20°18.6´w). plain immediately east of gravelven, east of mørkefjord station, sw germania land. named by the 1938–39 mørkefjord expedition. foraarsboplads was occasionally used by the 1906–08 danmark-ekspeditionen for this locality. gravstenene 71ø-70 (71°35.0´n 26°56.7´w). series of mountain summits up to 1800 m high on the ne side of nordvestfjord opposite the mouth of flyverfjord. the name originated during the 1931–34 treårsekspeditionen because the mountain tops resembled a row of gravestones, and was adopted at the suggestion of r. spärck. great claw – see the great claw. great cumbrae 71ø (71°56.6´n 25°05.7´w; map 5). upper branch of cantabræ, stauning alper. so named by the 1998 scottish moun taineering club expedition. great fault river 73ø (73°58.7´n 21°21.9´w). reference locality used by dunbar (1955) for one of lauge koch’s sample sites east of kap stosch, home forland. this is probably the river flowing in forkastningsdal (great fault valley), also called river 16, and officially known as blåelv. great fault valley – see forkastningsdal. great snow crest 72ø (72°19.7´n 25°38.4´w; map 5). snow ridge up to 2373 m high in ne nathorst land, nw of galenadal. it was the highest climb made by the 1970 st. andrews university expedition. (the great snow crest.) great white 70ø (70°48.8´n 26°10.9´w). marked summit 1645 m high on the north side of korridoren, milne land. climbed by the 2004 west lancashire scouts expedition via the glacier to the north of the summit, which they named great white glacier. green river 72ø (72°31.0´n 24°01.5´w). name used by the 1974 joint biological expedition for a minor stream on sw traill ø, west of karupelv, which drains into holm bugt. gregory cove 73ø (73°09.4´n 27°34.1´w). name used by n.e. odell (1939) during louise boyd’s 1933 expedition for the bay at the mouth of knækdalen (their gregory valley) in western strindberg land. it was named after. john walter gregory [1864–1932], a noted british geologist, who was drowned in 1932 in the gorge of the urumbamba, peru. gregory gletscher 73ø-540 (73°10.0´n 28°22.6´w; map 4). glacier flowing nne from petermann bjerg to the head of knækdalen, western strindberg land. named by j.m. wordie’s 1929 expedition as gregory glacier after j.w. gregory. see gregory cove. (gregorys gletscher.) gregory lake 73ø (c. 73°13´n 28°00´w). name occasionally used in louise boyd’s 1933 expedition reports for the traces of a former lake in knækdalen, western strindberg land, their gregory valley. gregory stream 73ø (73°11.6´n 27°39.8´w). name occasionally used in louise boyd’s 1933 expedition reports for the present knæk elven, western strindberg land, the river in knækdalen, their gregory valley. gregory valley 73ø (73°12.9´n 27°55.4´w). name originally used by j.m. wordie’s 1929 expedition for the valley in western frænkel land containing gregory gletscher and continuing northwards to the mysteriesøer. the name was subsequently applied by the louise boyd expedition of 1933 to the valley draining from gregory gletscher to the head of kejser franz joseph fjord, the present knækdalen. upper gregory valley and lower gregory valley were used for different sections. norwegian hunters used gregory dal in preference to the official name as late as the 1950s. gregorydalhytten 73ø (73°09.6´n 27°03.8´w). norwegian hunting hut built in april 1950 by arktisk næringsdrift east of the mouth of knækdalen in western strindberg land (also known as gregory dal – see gregory valley). the hut has also been known as bræhytten and knækelvhytten. greindalen 74ø (74°09.0´n 21°28.3´w). valley on south clavering ø, the lower part of the present granatdal. so named on nsiu maps of 1932 and 1937, because it has numerous branches (= grein) or tributary valleys. greindalsbreen 74ø (74°15.0´n 21°20.6´w). glacier on clavering ø, a branch of the present snemarken draining into greindalen. so named on the nsiu (1932a) map, but not distinguished on the nsiu map of lacmann (1937) where it is part of lars chris tensenfonna. it corresponds to the sw part of the present sne marken. the name derives from its proximity to greindalen. greipar 72ø-s4 73ø (72°00´–74°00´n 24°30'w). this name is one of several appearing in the icelandic sagas (e.g. hauksbók), which tornøe (1944) suggested might lie in east greenland. tornøe pro posed that the name, taken to mean ‘the space between fingers’ might have been applied to the fjord region between 72°–74°n, 183 rather than the disko region of west greenland as early authorities had proposed (rafn 1845). grejsdalen 73ø-647 (73°35.5´n 26°00.0´w; maps 2–4; see also fig. 74). major e–w valley in andrée land draining into kejser franz joseph fjord between eleonore bugt and kap weber. so named by ove simonsen during the 1931–34 treårsekspeditionen because of the rich vegetation and wildlife, after its danish namesake grejsdal, north of vejle, jylland. grejsdalshytten 73ø (73°28.5´n 25°02.9´w). norwegian hunting hut on the east side of the mouth of grejsdalen, andrée land, built for arktisk næringsdrift in march 1937. it was originally known as ragnhildshytta and has also been called eleonorebukta. grenen 81ø-74 (81°05.0´n 14°18.0´w; map 1, 4). eastern branch of flade isblink in north kronprins christian land. mapped and named by lauge koch during flights in 1933 during the 1931–34 treårsekspeditionen. the name probably derives from it being a branch of a larger glacier (grenen = the branch), though it may also have been named after the northernmost point of jylland, den mark. grey mound 73ø (73°32.2´n 25°49.0´w). mountain 1975 m high on the south side of grejsdalen, andrée land. climbed by the 2007 army boreal zenith expedition. grete gletscher 70ø-174 (70°38.3´n 22°04.8´w; map 4). glacier in south liverpool land draining west into gubbedal. so named by laurits bruhn during the 1931–34 treårsekspeditionen, together with hans gletscher and heksefjeldet, after the characters in the grimm brothers’ fairy tale ‘hänsel und gretel’ (hans og grete in danish; hansel and gretel in english). grifgletscher 74ø-373 (74°41.1´n 22°29.5´w). small corrie glacier on the ne side of grossglockner, thomas thomsen land, where glaciological studies were made by the 1948 leeds university expedition. the gryphon (or griffin), a mythical figure with an eagle’s head and wings and a lion’s body, is the emblem of leeds uni versity union, battle’s university home. (gryphon glacier.) grifhovedet 73ø-696 (73°17.3´n 26°12.5´w). mountain 1710 m high in south andrée land. named during lauge koch’s 1949–51 expeditions by john haller who made the first ascent in may 1950. the shape of the summit resembles the head of a bird. the griffin is associated in legend with basel, switzerland (where john haller was based). in january each year a festival is held in basel for ‘vogel gryff ’. grimm fjelde 76ø-341 (76°17.0´n 25°04.0´w; map 4; fig. 21). hilly region south of budolfi isstrøm, south dronning louise land. the name was given by the 1952–54 british north green land expedition, in association with the nearby features eventyr fjelde (eventyr = fairy tale) and h.c. andersen fjelde, and com me morates the german philologist and mythologist jacob ludwig carl grimm [1785–1865] and his brother wilhelm carl grimm [1786–1859]. the grimm brothers made a noted collection of folk tales transcribed from oral sources. grindøya 72ø (72°45.0´n 22°56.9´w). island in vega sund, equivalent to the present kista ø. the norwegian term ‘grind’ signifies a structure used to control the flow of water in a canal or channel. griper red 74ø-19 (74°32.1´n 18°52.3´w). anchorage off south sabine ø, west of germaniahavn. named by douglas clavering as griper roads (red = roads) after the griper, the ship of his 1823 voyage which had anchored here. the griper was a 180 ton sloopof-war, the same ship used by william parry on his 1st arctic voyage in 1819–20. gronau gletscher 69ø-41 (69°29.0´n 30°54.0´w). glacier in northern kong christian ix land, named by l.r. wager during his 1935–36 expedition as gronaus glacier after nearby gronau nunatakker. gronau nunatakker 69ø-34 (69°27.0´n 30°15.0´w; map 3). nuna tak area in north kong christian ix land discovered by wolfgang von gronau on 15 august 1931 during his flight from scoresby sund across the ice cap to sukkertoppen. mapped and named by lauge koch during flights in 1933 during the 1931–34 treårs ekspeditionen, and reported as undoubtedly the mountains seen by von gronau. wolfgang von gronau [1893–1977] was a pioneer aviator. (von gronau nunataks.) gross bjerg 73ø-285 (73°21.8´n 22°53.4´w). mountain about 1000 m high on the sw coast of gauss halvø. named during the 1931–34 treårsekspeditionen by gunnar säve-söderbergh as mt. gross, after walter robert gross [1903–74], a latvian vertebrate palaeontologist especially noted for studies of devonian fishes at the universities of frankfurt am main and tübingen. grosse kederbacher spids 71ø (71°52.8´n 25°37.4´w; map 5). moun tain on the west side of spærregletscher. named and climbed by the 1967 berchtesgadener expedition. it has also been called drillinge. grosse sirius pass 71ø (71°56.7´n 23°58.4´w; map 5). broad col between sirius gletscher and østre gletscher, between taget and øbjerg in the werner bjerge. the name is found on the maps of styger (1951), an account of climbing activities during lauge koch’s 1950 expedition. grosser sydney gletscher 71ø (71°57.5´n 25°41.1´w). name used by the 1967 berchtesgadener expedition for the tributary glacier west of spærregletscher more commonly known as castor glacier. it was named after sydney tinde at the head of the glacier. grosses becken – see skålen. grossglockner 74ø-68 (74°41.3´n 22°19.7´w; map 4). mountain massif 1300 m high nw of inner tyrolerdal, thomas thomsen land. named by karl koldewey’s 1869–70 expedition as gross glockner. discovered by julius payer in november 1869, it was named after the mountain of the same name in austria. see also pasterze and tyrolerfjord. (gross-glockner, mt. gross glockner.) grottedal 80ø-122 (80°22.7´n 21°39.3´w; map 4). valley north of centrumsø, kronprins christian land. named during operation groundhog 1960 for the presence of several limestone caves (davies & krinsley 1960). these were explored by a french speleological expedition in 1983. grottenfjeldet 80ø (80°04.8´n 22°37.4´w). name used by the 1983 french speleological expedition for a limestone cave in kronprins christian land at the corner between græselv and centrumsø. (grotte des quatre.) grouchs snack 75ø (75°59.4´n 19°53.5´w). southernmost skerry of the depotskærene, ene of trums ø. the name is used in den grønlandske lods (1968). gruberfjellet 74ø (74°18.0´n 21°04.7´w). mountain on central clav ering ø, west of skillegletscher. so named on the nsiu maps of lacmann (1937) after otto von gruber [1884–1942], who made significant contributions to the development of photogrammetry while working with carl zeiss, jena. grundtvigskirken 71ø-71 (71°06.9´n 25°57.9´w; map 3; fig. 44). mountain massif 1977 m high in renland, on the nw side of ø fjord, dominated by a granite tower bearing a remarkable resemblance (as seen from the east) to the tower of grundtvigskirken in copenhagen. a photograph of this peak appears in reports of the 1931–34 treårsekspeditionen (thorson 1937). due to poor topographic maps and an inaccurate description the name was positioned on some maps against a prominent peak 1882 m high 4 km to the sw of the original mountain (71°05´n 26°05´w; higgins 1986). the mountain was attempted by a party of four climbers from norway and sweden in 1998, who were successful with a second attempt in 1999; the climbing party gave it the name tsavagattaq (sometimes seen written as tasvagattaq). grunnvågen 72ø (72°51.8´n 22°00.0´w). bay on east geographical society ø, north of cambridge bugt. so named on the nsiu maps of lacmann (1937), because the bay is shallow (= grunn). gryden 73ø-87 (73°32.6´n 23°17.4´w). major depression in central gauss halvø at the head of vastidal. named by lauge koch’s 184 1929–30 expeditions (gryden = the bowl). grydepas 72ø-393 (72°02.9´n 23°20.3´w). minor pass at the head of medusagryde, north of kolledalen in northern scoresby land. named by hans kapp during lauge koch’s 1957–58 expeditions. grytvika 73ø (73°43.5´n 20°29.6´w). bay on the south side of knudshoved, on the east coast of hold with hope. so named on the nsiu (1932a) map for its cauldron-like shape. the name also appears in den grønlandske lods (1968). grænsebjerg 72ø-469 (71°59.3´n 26°44.8´w). mountain on the south side of grænsedal where it meets frederiksdal, nathorst land. named during lauge koch’s 1954–55 expeditions by hans zweifel. it is just south of latitude 72°n on modern maps. grænsedal 72ø-439 (72°01.0´n 26°52.1´w; map 4). e–w-trending valley running almost along the 72°n line of latitude. so named during the 1931–34 treårsekspeditionen by ove simonsen because latitude 72°n was the original south limit of surveying during the expedition (grænse = boundary, limit). grænsedalen 71ø (71°59´n 23°20´w). e–w-trending valley draining into antarctic havn, the present kolledalen. so named by hans stauber during the 1936–38 two-year expedition, because it was the north boundary (= grænse) of his working area. grænsedalen 74ø-353 (74°18.3´n 20°03.8´w). valley in southern wollaston forland. this name was originally used by frebold (1931), but not precisely delineated until the work of wolf maync and andreas vischer during the 1936–38 two-year expedition. grænseelv 74ø (74°28.1´n 20°29.8´w). minor river east of zacken berg forskningsstation draining into young sund. the name is used as a reference locality by visiting scientists (meltofte & thing 1996). (border river.) grænseryggen 74ø (72°29.8´w 19°34.4´w). name used by maync (1947) for the ridge north of gyldenspids in northern wollaston forland, which is bounded by a marked fault line. the name arose during work on lauge koch’s 1936–38 expeditions. (boundary ridge.) grænsesø 72ø-468 (72°00.8´n 27°16.0´w). lake in grænsedal, nat horst land, named by hans zweifel during lauge koch’s 1954–55 expeditions. græsdalen 72ø-175 (72°59´n 23°00´w). side valley to tværdal in central geographical society ø. named on the nsiu maps of lacmann (1937) in the form teigandalen for the clumps (= teigan) of grass. (græsdal). græselv 79ø-44 80ø-119 (80°03.6´n 23°00.0´w; maps 1, 4; fig. 24). valley in southern kronprins christian land draining north into centrumsø, with relatively luxuriant vegetation com pared to adjacent areas. named during operation groundhog 1960. (græselven, græselv river.) græstørvshytten 74ø (74°35.7´n 19°51.4´w). norwegian hunting hut built in august 1928 by the hird expedition on the west side of albrecht bugt, wollaston forland. the walls of the hut were sup ported by turf (= græstørv). the hut was more commonly known as sletta. grøfteelv [niinngarpik] 70ø-185 (70°31.2´n 22°23.5´w). river in south liverpool land draining west into hurry inlet. named during the 1931–34 treårsekspeditionen by laurits bruhn for the shape of the valley it occupies (grøft = ditch). grøndalsvatnet 74ø (74°14´n 20°37´w). lake in grønnedal, eastern clavering ø, so named on the nsiu maps of lacmann (1937). (gröndalsvatnet.) grønhorn 73ø-418 (73°58.5´n 27°53.1´w). nunatak in arnold escher land formed by greenish volcanic rocks. named during lauge koch’s 1951 expedition by hans r. katz. (grönhorn.) grønlands styrelse gletscher 69ø-37 (69°30.0´n 29°40.0´w). fig. 44. distinctive 1977 m high mountain on the north-west coast of renland, that has a remarkable resemblance to the tower of grundtvigskirken in copenhagen. when the 1931–34 treårsekspeditionen departed for east greenland only the tower of the copenhagen church had been completed, and a photograph of the mountain grundtvigskirken features as a landmark in thorsen’s (1937) popular account of the expedition. 185 glacier in the ice plateau region south of scoresby sund, which drains southwards. named by martin lindsay’s 1934 british trans-greenland expedition after grønlands styrelse, the danish administrative department responsible for greenland, subsequently the ministeriet for grønland (ministry for greenland). grønlænderhusene 74ø (74°15.1´n 19°47.0´w). a hut of this name is shown on jennov’s (1939) map sw of herschell bjerg, about 3 km east of blæsedalen, wollaston forland. it was built by nanok in july 1930, washed away by a storm in 1931, and replaced by a new hut in 1935. the original hut was built on the site of an inuit house. (grønlænderhuset.) grønlænderhuset 74ø (74°15.9´n 19°22.9´w). name used by hunters of østgrønlandske fangstkompagni for the hunting hut at kap borlase warren in se wollaston forland. an old inuit house had been used as a hunting hut by severin liavaag’s 1908–09 expedition, which had called it borganes. a danish station built on this site in 1922 (valdemarshaab) replaced the hut, but was taken down in 1923, and the inuit house was again taken into use. see also borganes. grønne nunatak 78ø-27 (78°29.0´n 23°00.0´w; map 4). nunatak in the garde nunatakker group, south of zachariae isstrøm, de scribed by eigil knuth (1942) as the largest nunatak (grøn = green). the name is misplaced on some maps to one of the smaller nunataks to the nw. grønnebjerge 72ø-231 (72°40.0´n 23°24.7´w). mountain range up to 950 m high on ne traill ø, south of rold bjerge. so named by desmond t. donovan during lauge koch’s 1949–50 expedition for the greenish colour of the rocks. grønnedal 74ø-110 (74°13.5´n 20°26.4´w; map 4). valley on eastern clavering ø. the name was reported by seidenfaden (1931) as in common use by danish hunters, and was subsequently adopted in scientific reports. a sirius hut built between 1950 and 1960 about 10 km up the valley (74°13.5´n 20°31.1´w) is also known by the name grønnedal (p.s. mikkelsen 1994, 2008). (grönnedal, grønnedal valley, green valley, gröndalen.) grønnedalshytten 74ø (74°13.8´n 20°31.7´w). danish hunting hut built in april 1947 by nanok about 6–8 km up grønnedal, clav ering ø. it was destroyed by wind in the spring of 1950, and re placed by the sirius hut known as grønnedal (see above). grønnesø 71ø-326 (71°59.7´n 28°57.6´w). lake in charcot land. the name was approved in 1961 at the suggestion of ulrik røen, and records the green colour of the water. grønningen 74ø-370 (74°48.0´n 21°47.9´w). valley in th. thom sen land draining into svejstrup dal. it was named by the 1948 leeds university expedition because of the plentiful grass and vegetation. (green valley.) grønsø 72ø-377 (72°00.3´n 23°41.9´w). small lake on the north side of kolledalen, north scoresby land. named by hans kapp during lauge koch’s 1957–58 expedition, for the colour of the lake. gråbeinryggen 73ø (73°35.0´n 21°12.2´w). ridge in the southern tågefjeldene, hold with hope. so named on an nsiu map (1932a; fig. 13), and probably derived from the norwegian dialect word for a wolf (= gråbein). gråfjeld 79ø-32 (79°59.5´n 20°18.3´w; map 4). mountain on the west side of dijmphna sund, south of the mouth of rivieradal. the name was suggested by the place name committee in 1960 as a replacement for a proposal by john haller. it records the colour of the rocks. gråfjellet 73ø (73°08.5´n 23°30.5´w). mountain 1099 m high on eastern ymer ø, south of dusén fjord. the name appears in this form on an nsiu map (1932a), and appears to be identical with the present teglbjerg. graah bugt 71ø (72°01.9´n 28°30.9´w). name used by helge g. backlund during the 1931–34 treårsekspeditionen for the innermost section of nordvestfjord in front of f. graae gletscher (occasionally incorrectly referred to as graah gletscher). the name arose because of the assumption that the glacier had been named after the danish naval officer wilhelm august graah [1793–1863]. see kap graah. graah gletscher 72ø (72°06.8´n 28°42.3´w). name used on some of lauge koch’s geological map compilations (e.g. koch & haller 1971) for f. graae gletscher. see also graah bugt. gråhorn 73ø-673 (73°37.5´n 26°33.2´w). mountain in west central andrée land, on the south side of gneisdal. named during lauge koch’s 1949–51 expeditions by john haller, for the grey colour of the rocks. gråhoved 72ø-466 (72°54.1´n 29°03.1´w). nunatak on the west side of upper nordenskiöld gletscher, west of shackleton bjerg. it was named by john haller following explorations during lauge koch’s 1953 expedition, presumably for the shape and colour of the nunatak. gråklint 71ø-445 (71°45.9´n 22°56.5´w). characteristic greycoloured cliff ne of the solfaldsdal delta on the north side of fleming fjord. named by lars b. clemmensen during field work with ggu’s 1976 expedition. gråkløft 70ø-39 (70°44.1´n 25°18.8´w). ravine on the coast of eastern milne land between charcot havn and kap leslie. named during the 1931–34 treårsekspeditionen by hermann aldinger as graue schlucht (= grey ravine) for the colour of the rocks. (gray ravine). gråkollen 73ø (73°11.1´n 25°58.4´w). norwegian hunting hut built in july 1947 for arktisk næringsdrift on the north coast of suess land. it is commonly known as polarheimen. gråvæggen 72ø-391 (72°03.5´n 23°18.2´w). mountain wall nw of antarctic havn, north scoresby land, formed by a grey gabbrodiorite intrusion. named by hans kapp during lauge koch’s 1957–58 expeditions. guardian of korridoren 70ø (70°48.7´n 25°58.0´w). summit about 1490 m high on the south side of korridoren, milne land, that is a conspicuous feature when ascending korridoren from the east. named by the 2004 west lancashire scouts expedition. gubbedal 70ø-178 (70°37.5´n 22°17.3´w). valley in liverpool land draining west to hurry inlet (gubbe = old man). so named during the 1931–34 treårsekspeditionen by laurits bruhn, possibly after the rumanian scientist, constantin dumbrava, who had built a house at the mouth of the valley in 1934 – see also dum brava. dombravadal has also been used. gudenelv 72ø-88 (72°28.7´n 23°04.1´w; map 4). river on traill ø flowing se into mountnorris fjord, named during the 1931–34 treårsekspeditionen by ove simonsen after the danish river gudenå in jylland. guglia della norsketinde 72ø (c. 72°08´n 25°03´w). peak 2400 m high in the northern stauning alper, in the vicinity of norske tinden. it was named and climbed by g. dionisi’s 1982 expedition. guiden 69ø-38 (69°07.0´n 29°47.0´w). nunatak 2926 m high on the east side of christian iv gletscher. it was used by sledging parties as a steering mark, and was originally termed the guider. guldhorn 73ø-383 (73°40.7´n 25°34.2´w). mountain 1851 m high in eastern andrée land, north of grejsdalen. so named during lauge koch’s 1948–50 expeditions by erdhardt fränkl because the summit was formed by yellow quartzite (guld = gold). guldtinderne 72ø-292 (72°55.7´n 28°28.2´w). two mountain sum mits 2400 m high in southern goodenough land. named by john haller following explorations during lauge koch’s 1953 expedition, presumably for the colour of the rocks forming the sum mit. (guldzinnen.) guldtoppen 74ø (c. 74°16´n 19°23´). name reported used by the 1908–09 floren expedition for a hill in the vicinity of kap borlase warren (brandal 1930). exact position uncertain. gule horn 71ø-348 (71°20.8´n 22°42.9´w). mountain 975 m high in eastern jameson land, west of inner carlsberg fjord. named in 186 geological reports during lauge koch’s 1958 expedition by john h. callomon, for the colour of the rocks (gule = yellow). gulelv 73ø-48 (73°56.2´n 21°13.6´w). river in home forland, northern hold with hope draining north. named by lauge koch’s 1929–30 expeditions in the form yellow river, probably for the colour of the triassic rocks. it has also been referred to as river 19 (koch 1931). (gula, gulaelv.) gulfjelde 70ø-190 (70°30.8´n 22°10.5´w). mountains of yellow sandstone about 300 m high on the west side of rosenvinge bugt, southern liverpool land. named during the 1924–25 colonisation expedition. (gule fjælde, yellow fjeld, montagne jaune.) gully–lang col 72ø (72°03.6´n 24°55.9´w; map 5). pass in the north ern stauning alper between the head of gully gletscher and storgletscher (formerly langgletscher). gullygletscher 72ø-79a (72°06.3´n 25°16.4´w; maps 4, 5). glacier occupying a deep and spectacular gully in the stauning alper. the name originated from j.m. wordie’s 1929 expedition, and appears to have been used originally as an alternative name for sefström gletscher, with which the present gullygletscher merges to almost block alpefjord. wegmann (1935) designated the two glaciers gully-gl.1 and gully-gl.2. gulmann sund 73ø-121 (73°53.9´n 20°14.9´w). sound between jackson ø and home forland, ne hold with hope. the name was in use by hunters of østgrønlandske fangstkompagni from about 1923, and is said to originate with gustav thostrup, captain of the teddy in 1922. it was named after christian gulmann [1869–1934], journalist and editor of the danish newspaper berlingske tidende from 1912. (gulmans sound.) gultop 77ø-136 (77°05.4´n 23°56.1´w; map 4; fig. 21). mountain in northern dronning louise land at the nw edge of ad astra iskappe. so named by the 1952–54 british north greenland expedition because of the yellow quartzite forming its summit. (guletop.) gultop gletscher 77ø-135 (77°04.0´n 24°01.0´w; fig. 21). small glacier flowing from ad astra iskappe, near gultop, to the snout of admiralty gletscher, north dronning louise land. named by the 1952–54 british north greenland expedition. gundahl knold 76ø-123 (76°42.1´n 23°02.3´w; fig. 21). isolated hill in eastern dronning louise land at the front of borgjøkel. named by j.p. koch’s 1912–13 expedition as gundahls knold, after jens gundahl knudsen [1876–1948]. he was the carpenter on the 1906–08 danmark-ekspeditionen, where he built the expedition house, meteorological station and expedition sledges. he also worked in west greenland at a copper mine near ivigtut from 1910 to 1912. gunnar andersson land 73ø-23 (73°20.5´n 24°22.5´w; maps 3, 4). north part of ymer ø, north of dusén fjord. in 1929 lauge koch followed up the reports by his greenlandic hunters that dusén fjord was longer than it was thought to be, and found that it almost divided ymer ø into two parts. he named the northern part after carl filip gunnar andersson [1865–1928], a swedish geographer who was koch’s father-in-law, and also editor of the swedish journal ymer for 27 years. the day of koch’s observations was just one year after andersson’s death (koch 1930b). (gunnar anderssons land.) gunnar hornsletta 72ø (72°58.0´n 22°21.5´w). low-lying area on the ne side of geographical society ø. used only on nsiu maps (lacmann 1937), the name commemorates gunnar horn [1894– 1946], an nsiu geologist who participated in several expeditions to svalbard and franz joseph land between 1924 and 1930, and is most noted as leader of the expedition that discoved the remains of andrée’s balloon expedition on kvitøya in 1930. he visited se greenland on a 1932 nsiu expedition. gunnbjörnfjellet 73ø (73°22.3´n 22°57.5´w). mountain on the south side of gauss halvø, corresponding to the west end of the present hjelmbjergene. so named on an nsiu map (1932a) after gunn björn ulfsson, noted for the discovery of skerries off se greenland which bear his name. the name appears as mt gunnbjørn, and in correctly as mt grimbjørn, on plates 1 and 2 of säve-söderbergh (1933). (gunnbjørns bjerg.) gunner andersen sø 76ø-179 (76°06.5´n 20°14.7´w). lake in ad. s. jensen land. the name was proposed by the place name com mittee as a substitute for the rejected nanok name frieda sø. it commemorates gunner andersen, a nanok telegraphist who died of exhaustion when overtaken by bad weather during a sledge journey in april 1933. gunnsteinfjellet 73ø (73°28.2´n 22°04.8´w). mountain north of gustav dal in the giesecke bjerge, corresponding to the east sum mit of troels-lund bjerg. used only on the nsiu (1932a) map. (gunnsteinsbjærg.) guntherbreen 74ø (74°17.7´n 21°17.7´w). glacier on central clav ering ø, draining west. so named on nsiu maps of lacmann (1937), after gunther, who married brunhilde in the german epic poem from c. 1200, the nibelungenlied. gunvor bjerg 73ø-568 (73°41.1´n 24°36.7´w). mountain 1231 m high in strindberg land. named during the 1931–34 treårseks peditionen in the form mt gunvor, apparently after paul gelting’s wife. gelting visited the area in november 1931. (gunvørs bjerg, gunvors bjerg.) gurreholm 71ø-159 (71°14.7´n 24°35.0´w; maps 3, 5). danish scientific station in western jameson land near nordostbugt, built in 1937. it was named after a summerhouse of that name belonging to the owner of the bulldog shopping chain, who had made a large contribution to lauge koch’s expedition finances. this scientific station was originally planned for a site in the interior of fleming fjord, but ice conditions in 1937 made access impossible. lauge koch’s expeditions used the station in 1937–38 when icelandic ponies were stationed here and employed for transport in the wide expanses of jameson land. it was occasionally occupied during the war years when it went under the code name bluie east 3. post-war, it was used occasionally by danes and greenlanders from scores bysund, and nordisk mineselskab also made use of the building during mineral prospecting in the 1970s. fuchs (1984) mistakenly refered to it as alfred wegener’s eastern station, but this was sited some distance to the south (see tyskit nunat). inuit ruins in this vicinity have been referred to under the greenlandic name itti kajik. gurreholm bjerge 71ø-161 (71°42.7´n 24°05.0´w; maps 4, 5). mountain range up to 1360 m high in western jameson land, about 30 km nnw of gurreholm station on the west side of schuchert dal. the mountains were named by hans stauber during lauge koch’s 1936–38 expeditions. (gurreholmberge, gurreholm-bjerge.) gurreholm dal 71ø-160 (71°26.0´n 24°43.5´w; maps 4, 5). valley nnw of gurreholm station in western jameson land. named during lauge koch’s 1936–38 expeditions by hans stauber. (gurreholmtal.) gurreholm slugt 71ø-371 (71°24.5´n 24°41.5´w; map 5). ravine at the mouth of gurreholm dal, western jameson land, draining south into nordostbugt. named by the 1962 oxford university expedition. (gurreholms gorge.) gurreholms elv 71ø (71°26.0´n 24°43.5´w). name used by the 1962 oxford university expedition (sugden & john 1965) for the river flowing in gurreholm dal. gustav dal 73ø-344 (73°27.5´n 22°04.6´w). valley in the giesecke bjerge, eastern gauss halvø, draining eastwards. named by wolf maync and andreas vischer during lauge koch’s 1936–38 expeditions after the greenlander who assisted them in the summer of 1938. gustav thostrup bjerg 79ø-31 (79°56.9´n 19°34.3´w). mountain c. 1200 m high on western hovgaard ø. named during lauge koch’s 1956–58 expeditions by john haller after gustav gustavsen thostrup [1877–1955], 2nd mate and surveyor on the 1906–08 187 danmark-ekspeditionen. he subsequently took part in several voyages to east greenland as pilot or ice-pilot, including that of the dagny in 1919 and the teddy in 1921 and 1922. guthrie bjerg 69ø-76 (69°43.2´n 23°53.0´w). mountain in henry land, on the northern blosseville kyst. named by malcolm sles ser’s 1969 expedition after the old tumble-down quarter of the small royal burgh of brechin in the tayside region of scotland, where one of the expedition members lived. the mountain was climbed on 8 august, and described as comprising ‘tumble-down’ rocks. (mt guthrie.) gyldenspids 74ø-345 (74°29.0´n 19°37.2´w). mountain about 660 m high in wollaston forland. so named by wolf maync and andreas vischer during lauge koch’s 1936–38 expeditions, because the pyramid-shaped summit is formed of golden-yellow sedimentary rocks (maync 1947). (goldene spitze.) gylfeelv 72ø-194 (72°11.1´n 24°04.9´w; map 5). river in northern scoresby land, sw of mestersvig, draining from the sw flank of domkirken into store blydal. named by prospecting teams associated with lauge koch’s 1948–49 expeditions. gylfe was the swe dish saga king who gave gefion all the land she could plough in a day, which is said to have resulted in the separation of the island of sjælland from sweden. gänsepingo 72ø (72°36.1´n 23°41.8´w). name used by fritz müller during the 1954–55 lauge koch expeditions for a pingo in a side valley north of karupelv, traill ø. it was named for the geese (= gänse) in the area. gaase dal 70ø (70°46.7´n 22°46.7´w). name used by rosenkrantz (1934) for the valley in east jameson land in which gåseelv flows. gåsebugt 72ø-159 (72°12.4´n 22°10.1´w; map 4). bay on se traill ø between kap young and kap moorsom. named during lauge koch’s 1936–38 expeditions by hans peter schaub for the presence of geese. (gaasebugt.) gåsedal 74ø-344 (74°30.0´n 19°12.5´w). valley in ne wollaston forland, so named during lauge koch’s 1936–38 expeditions by wolf maync and andreas vischer for the geese. gåsedal 74ø (74°28.3´n 20°28.6´w). valley east of zackenberg forskningsstation. the name has been used by visiting scientists. gåseelv 74ø (74°28.3´n 20°28.6´w). river east of zackenberg forskningsstation. the name has been used by visiting scientists. gåseelv 70ø-108 (70°46.7´n 22°46.7´w; map 4). river on the west side of hurry inlet entering the fjord north of constable pynt. named during lauge koch’s 1926–27 expeditions by alfred rosen krantz, although the name first appeared in harris (1931) without a precise location. the present position is that of rosen krantz (1934), who used the form gaase elv. it was named for the numerous barnacle geese. the same name was used in error for the present primulaelv by roberts (1935). (gaaseelv, r. gaase elv.) gåseelv 76ø-305 (76°58.3´n 20°10.3´w). river in western ger mania land, flowing west through gåsesøen and into the south end of sælsøen. named by the 1938–39 mørkefjord expedition. (gaaseelv.) gåsefjeldet 76ø-354 (76°23.8´n 20°55.9´w). mountain north of ålborghus on godfred hansen ø, western dove bugt. the name was reported by hans meltofte as in regular use by personnel at danmarkshavn weather station because of the large barnaclegoose colony, and is said to have first been used by danish hunters (jennov 1963). gåsefjord [nertivit kangersivat] 70ø-17 (70°10.0´n 27°15.0´w; maps 3, 4). large e–w-trending fjord south of gåseland named by carl ryder’s 1891–92 expedition as gaasefjord (fig. 7). barnacle geese and pink-footed geese are very common throughout the scoresby sund region. sydfjorden was used for the same fjord in ragnvald knudsen’s diaries of the expedition, and taagefjord in a report by nikolaj hartz. (gaasefjorden, gaase fiord, gaase fjord, gänse-fjord.) gåseflade 70ø-385 (70°10.0´n 28°41.3´w). part of vindblæsedal south of faxe sø in western gåseland. named during lauge koch’s 1958 expedition by eduard wenk, for the numerous geese. (gaase flade.) gåsegletscher 70ø-387 (70°02.8´n 28°38.4´w; maps 3, 4). glacier draining into the head of gåsefjord. named during lauge koch’s 1958 expedition by eduard wenk. (gaasegletscher.) gåseholm – see gåseholmhytten. gåseholmhytten 75ø-94 (75°58.5´n 21°52.0´w). danish hunting hut at the west end of bessel fjord, built by nanok in august 1938 at the east end of a narrow gravel spit (gåseholm). named for the geese, which breed commonly in the region. the hut was still standing in 1990, but is a ruin. (gåseholmshytten, gaaseholmhytten). gåsehytten 72ø (72°50.8´n 22°56.8´w). name sometimes used for the norwegian hunting station built by arktisk næringsdrift in 1929 in geographical society ø on the north side of vega sund, about 5 km nw of gåseøen. it is usually known as sverresborg. gåseland 70ø-15 (70°15.0´n 28°00.0´w; maps 3, 4). large peninsula or landmass between gåsefjord to the south and fønfjord and vestfjord to the north. named by carl ryder’s 1891–92 expedition in the form gaaseland for the geese. see also gåsefjord. vestlandet has also been used. (gaaselandet, gaase land, gåslandet, gänseland.) gåselien 74ø-306 (74°05.5´n 21°16.4´w). slope on the sw side of østhavn, close to eskimonæs station, clavering ø. the name originated from the wintering party at eskimonæs during the 1931–34 treårsekspeditionen and was given for the grazing geese. gaasepas 73ø (c. 73°44´n 20°27´w). name used by gelting (1937) for a locality near knudshoved, hold with hope, exact locality uncertain. it may have have been a danish hunters name. gåsepynt [ujuaakajiip nunaata akia] 70ø-69 (70°22.0´n 26°18.0´w; maps 3, 4). eastern cape of gåseland, named by carl ryder’s 1891–92 expedition as gaasepynt. gulløv (1991) also uses gåselandspynten (j. løve, personal communication 2010). see also gåsefjord. gåsereden – see kalkdalen. gåseslette 81ø (81°08.7´n 13°08.3´w). plain in eastern kilen, kron prins christian land where large flocks of barnacle geese congregate. the name is found on a coloured geological map of kilen printed in 1991 (pedersen 1991). gåsesø 72ø-201 (72°14.4´n 23°53.9´w). small lake west of noret, close to mestersvig airfield. named by prospecting teams associated with lauge koch’s 1948–49 expeditions. gåsesø 74ø (74°28.7′n 20°29.4′w). small lake east of zackenberg forskningsstation. the name has been used as a reference locality by visiting scientists. gaasesøen 70ø (70°27.6´n 26°16.5´w). lake west of blåbærhøj on danmark ø, limited to the west and north by steep slopes. so named during carl ryder’s 1891–92 expedition because geese were seen nesting here. the name is used in the report by hartz (1895). gåsesøen 77ø-22b (76°58.6´n 20°08.8´w). small lake east of the se end of sælsøen. named by the 1906–08 danmark-ekspeditionen as gaasesöen, because numerous traces of barnacle geese were seen here in may 1907. due to inaccurate topographic maps, the position of the lake was ‘officially’ placed north of latitude 77°n, but on modern topographic maps from 1989 the lake is just south of latitude 77°n. gåseøen 72ø-75 (72°47.9´n 22°54.1´w). flat island in vega sund, north of the scott keltie øer. named by nsiu in 1929 as gåsøya, because of the abundant signs of breeding geese and eider ducks. it was considered by a.k. orvin and b. lynge as the best bird terrain they had seen. later nsiu maps used fugleøya for the same location. (gaaseø.) gåshamrane 73ø (73°04.3´n 23°04.2´w). western cliff of robertson ø at the eastern end of sofia sund. used on the nsiu maps of lac mann (1937), the name records a breeding locality for geese. gåsneshuset 74ø (74°29.4´n 18°59.6´w). norwegian hunting hut 188 built by arktisk næringsdrift in 1928 close to kap wynn, the eastern cape of wollaston forland. named after severin gaasnes lia vaag [1879–1909], leader of the 1908–09 expedition, skipper and part-owner of the floren, who was drowned during a bear hunt between kap wynn and hvalrosø in may 1909. no trace of the hut remains. (gaasneshuset, gåsenes.) gåssjøen 73ø (73°59.4´n 23°47.0´w). lake in ole rømer land, the sw end of the present krumme langsø. so named by sigurd skaun and harald welde in 1932 because of the large number of geese along its banks. (gaasesjö.) h h.a. jensen bjerg 77ø-101 (77°10.6´n 23°43.6´w; map 4). mountain in northern dronning louise land, with a cairn on the summit said to have been built by the 1906–08 danmark-eks peditionen (hamilton et al. 1956). the name was given by the 1952–54 british north greenland expedition for hans a. jensen, danish surveyor on the expedition who died in a fall on 2 april 1953 near kap niels. a memorial inscription is carved on a boulder at danmarkshavn where his body was taken for shipment to denmark. h.c. andersen fjelde 76ø-342 (76°19.5´n 24°12.8´w; map 4; fig. 21). hilly region south of the lower part of budolfi isstrøm, southern dronning louise land. the name was given by the 1952–54 british north greenland expedition, in association with the names grimm fjelde and eventyrfjelde, for the danish author and poet hans christian andersen [1805–75]. h.g. backlund fjord 77ø-143 (77°30.5´n 20°24.9´w; map 4). small fjord in the inner part of skærfjorden, named by john haller following explorations during lauge koch’s 1956–58 expeditions. it was said to have been one of the last unexplored fjords in caledonian crystalline rocks, and thus suitably named after the noted swedish petrographer. see also backlund bjerg. h.l. jensens hus 74ø (c. 74°32´n 18°48´w). hut built at germania havn, sabine ø, by østgrønlandske fangstkompagni in the sum mer of 1922, and named after hans ludvig jensen [1874–1948], also known as ‘grønlands-hans’. one of the founders of the company, he had previously participated in the 1906–08 danmarkekspeditionen. the hut was taken down in 1923. see also germaniahavn. ha-ha-hytta 73ø (73°21.0´n 26°28.0´w). name occasionally used for the norwegian hunting hut in renbugten, southern andrée land, usually known as renbugthytten or reinsbukta. haakonshytta – see håkonshytta. hagar bjerg 72ø-420 (72°53.7´n 27°49.0´w; map 4). mountain 2470 m high in southern goodenough land, named during the 1931–34 treårsekspeditionen by eugène wegmann as mt. hagar after a swiss mountain of similar name. hagar massif has also been used in a broader sense for the mountain range including hagar bjerg. j.f.b. mountain has also been used. hagen [risip qaarusaa] 70ø-217 (70°39.3´n 21°36.5´w). peninsula on the east coast of liverpool land. so named by laurits bruhn during the 1931–34 treårsekspeditionen for its shape on the map relative to the nearby features gabet and snuden (hagen = the chin). hagen ø 77ø-49a (77°57.0´n 19°46.0´w; map 4). island in the southern part of jøkelbugten. named by the 1906–08 danmarkekspeditionen as hagens ø, after niels peter høeg hagen [1877– 1907], a danish army officer who was the expedition cartographer. hagen camped here for two days in the autumn of 1906. he was one of the three men of the 1906–08 danmark-ekspeditionen who died during the return from a sledge journey to independence fjord. the island has occasionally been called observationsø. (hagens ö.) hagenbreen 74ø (74°23.5´n 21°14.7´w). glacier on north clavering ø, draining westwards. so named on nsiu maps of lacmann (1937), after hagen, who killed the hero sigfred in the german epic poem from c. 1200, the nibelungenlied. hagenpasset 72ø-n140 (72°57.3´n 23°54.3´w). col or pass in west geographical society ø. so named on nsiu maps of lacmann (1937) after asbjørn hagen [b. 1912], a norwegian who participated as botanist in the 1933 nsiu expedition to east greenland. hahnenkamm 72ø (c. 72°12´n 25°10´w). mountain in the vikinge bræ region of the stauning alper. it was attempted unsuccessfully by hermann huber’s 1968 expedition, and climbed in 1970 by a party led by wolfgang weinzierl. the name was given for the austrian mountain near kitzbühel, site of the classic downhill ski run. exact location uncertain. (coxcomb.) hakkemandstoppene 77ø-27 (77°41.8´n 20°23.2´w). mountain in stormlandet, north of penthievre fjord, so named by the 1906–08 danmark-ekspeditionen. ‘hakkemand’ is a constructed danish word, from hakke (= chop) and mand (= man), perhaps inspired by hakon jarner’s nickname ‘hakke’ known to have been used by the family (j. løve, personal communication 2009). hall bredning 70ø-72 71ø-123 (70°54´n 25°45´w; maps 3, 4). very wide inlet between jameson land and milne land, named by william scoresby jr. in 1822 as hall’s inlet out of respect to basil hall [1788–1844]. a captain in the royal navy, hall published journals of several of his voyages. (hall fjord, halls inlet, hatts fjord.) hallebjergene 74ø-124 (74°15.0´n 21°51.0´w; map 4). range of mountains up to 1200 m high on west clavering ø, named by lauge koch’s 1929–30 expeditions in the form halle mtns. named after thore gustav halle [1884–1964], a professor at the university of stockholm who had worked on plant material from the expedition. hallehytta 73ø (73°33.6´n 22°44.0´w). norwegian hunting hut on the north side of moskusoksefjord, at the mouth of prospektdal in southern hudson land. it was built by finn devold’s expedition in september 1929 and was named after thor halle, a norwegian hunter with arktisk næringsdrift from 1929 to 1931. (halle-hytta, halle.) hallvardvatnet 72ø (72°53.3´n 22°08.7´w). lake on eastern geographical society ø. used only on nsiu maps (lacmann 1937), and so named after hallvard ophus devold [1898–1957], a norwegian telegraphist, meteorologist and hunter. see also devold hytta. halsneshytta 75ø (75°24.8´n 21°11.3´w). norwegian hunting hut on the north side of ardencaple fjord, built in august 1932 for john giæver’s expedition. now a ruin (1988). it is also known as holms nes and berglann. halveøen 74ø (74°27.0´n 20°26.4´w). peninsula on the coast of zackenberg bugt, south of zackenberg forskningsstation. the name has been used as a reference locality in reports by visiting scientists. ham-gletscher 72ø (72°00.9´n 24°07.4´w; map 5). central of three small glaciers between vestre gletscher and mellem gletscher in the northern werner bjerge. the name was used by styger (1951) in a report on a climbing excursion during lauge koch’s 1950 expedition, and with two of his other names (sem-gletscher and joffert-gletscher) commemorates the sons of noah. hamberg gletscher 73ø-571 (73°33.0´n 29°38.0´w; maps 2–4). major glacier west of louise boyd land, which swings ne to join gerard de geer gletscher. the name first appeared on the 1932 1:1 million scale geodetic institute map drawn by lauge koch during the 1931–34 treårsekspeditionen. it was given for axel hamberg [1863–1933], a swedish mineralogist and geographer, who was professor of geography at the university of uppsala from 1907 to 1928. he took part in several arctic expeditions. hamlet bjerg 72ø-291 (72°50.9´n 28°34.2´w). mountain 2390 m high in southern goodenough land. named during lauge koch’s 1953 expedition by john haller, whose party made the first ascent 189 on 1 august that year. it was named after hamlet, prince of den mark in shakespeare’s play of the same name. hammar ø 72ø-38 (72°32.5´n 24°38.3´w). island off the east coast of lyell land. named by a.g. nathorst’s 1899 expedition as hammars ö after josef hammar [1868–1927], an army doctor who was surgeon on nathorst’s expedition and also made rich ethnographical collections. (hammar island, hammer island, hamme r öya.) hammarskjöld brae 72ø (72°02.1´n 27°59.5´w). glacier in sw nathorst land, draining south to nordvestfjord, the present uni versitets gletscher. the name was introduced by geoffrey halli day during the 1961 leicester university expedition, and com me morates dag hammarskjöld, a swedish diplomat who died in an aeroplane crash on 18 september 1961. hammeren 73ø-658 (73°23.7´n 24°44.1´w; map 4). mountain 1427 m high in gunnar andersson land, northern ymer ø. named by th. johansen during the 1931–34 treårsekspeditionen, probably after the north point of the danish island of bornholm (ham meren = the hammer). hammeren 74ø-349 (74°22.7´n 19°52.8´w). mountain 1008 m high in central wollaston forland. the name was proposed by the place name committee in 1939, and was given for the north point of the danish island of bornholm. hammeren 78ø-23 (78°16.6´n 19°34.5´w; maps 1, 4). large island in jøkelbugten. named by the 1938–39 mørkefjord expedition together with stigbøjlen and ambolten for a supposed resemblance in shape to bones in the ear (hammeren = the hammer). (ham merøen.) hamna hytten – see havna. hamspitze 72ø (72°00.4´n 24°09.1´w). mountain about 1300 m high at the head of ham-gletscher, northern werner bjerge. the name was used in styger’s (1951) account of climbing activities during lauge koch’s 1950 expedition. see also ham-gletscher. hans gletscher 70ø-172 (70°40.7´n 22°06.5´w; map 4). glacier in southern liverpool land draining west into nøkkedal. named during the 1931–34 treårsekspeditionen by laurits bruhn, together with heksefjeldet and grete gletscher, after the characters in the grimm brothers’ fairy tale ‘hänsel und gretel’ (hans og grete in danish; hansel and gretel in english). hansa bugt 74ø-45 (74°37.5´n 18°47.1´w; map 4). enclosed bay on ne sabine ø. named by karl koldewey’s 1869–70 expedition as hansa bai, after the second ship of his expedition which was trapped in the ice and sank off liverpool land in east greenland. the crew drifted south with the ice and came ashore in frederiks dal, sw greenland. the hansa was a 77-ton prussian schooner, built in 1864 as the fulton, and renamed for the expedition. a hunting hut was built on the south side of the bay in 1928 by the norwegian hird expedition, who called the bay and hut ingridhavn. the hut disappeared during the war years. in 1942–43 a german meteorological station, comprising two huts (alte hütte and neue hütte), operated from the bay until bombed by the us air force on 25 may 1943. (hansa bay, hansabugten, ingridhavn.) hansa bugthytten 74ø (74°38.0´n 18°44.1´w). danish hunting hut on a small skerry off hansa bugt, sabine ø, built by nanok in august 1948. (hansabugten, hansa bugt hytten). hansabugthuset – see ingridhavn. hanseeraq fjord 80ø-35 (80°18.0´n 16°19.5´w; maps 1, 4). small fjord on the east coast of holm land. named by eigil knuth during his 1938–39 expedition after the greenlander johannes christian hansen (hansêrak) [1837–1911] of sydprøven, who had been a member of gustav holm’s 1883–85 konebaadsekspedition to ammassalik. (hansêraqs fjord, hansêraks fjord, hansêraq fjord.) hansen havn 74ø-162 (74°09.9´n 22°18.5´w). small bay north of jordan hill, named by lauge koch’s 1929–30 expeditions after p.m.j. hansen, first mate on the godthaab on the 1929 voyage. (hansen hafen, hansen harbour, hansenfjorden, hansenfjorden.) hansen havnhytten 74ø (c. 74°11´n 22°13´w w). norwegian hunting hut built in 1935 for arktisk næringsdrift about 3 km ne of hansen havn. now disappeared. it was also known as blåræven. harald grieg fjeld 73ø-582 (74°00.0´n 27°44.2´w). mountain in eastern arnold escher land, named in 1931 by arne høygaard and martin mehren as harald griegs fjell after harald grieg [1894– 1972], a norwegian publisher. the original usage was for a broader region of nunataks between skråbræ and the present mountain, somewhat larger than the present arnold escher land. the mountain was climbed by a party led by hans r. katz in august 1951. haraldsborg 75ø (75°15.1´n 18°49.4´w). danish hunting hut on the west coast of shannon about 10 km south of kap copeland, built for nanok in september 1948. it was named after harald mik kelsen who helped build the hut. it is also known as kap copeland hytten or copelandshytten. harder bjerg 73ø-83 (73°25.4´n 22°51.6´w). mountain 1675 m high on gauss halvø, named by lauge koch’s 1929–30 expeditions in the form mt. harder. probably named after the danish geologist p.j. harder [1878–1931]. harebjerg 74ø-51 (74°34.1´n 19°00.4´w). mountain 575 m high on sabine ø, named by karl koldewey’s 1869–70 expedition as hasen berg, because arctic hares were seen frequently here during the expedition (hase = hare; fig. 45). (mt hazenberg.) haredal 74ø-105 (74°20.7´n 19°16.4´w). valley in east wollaston forland south of clark bjerg. the name was reported by seiden faden (1931) as a danish hunters name, but is probably identical with the haredalen of severin liavaag’s 1908–09 expedition (brandal 1930). (hare valley). haredalen 73ø-600 (73°21.6´n 27°12.2´w; map 4). valley on the ne side of frænkel land, so named during the 1931–34 treårseks peditionen by gunnar thorson for the many hares seen here. haredalshytten 74ø (74°18.1´n 19°18.9´w). norwegian hunting hut on the east side of wollaston forland, built by the møre expedition in july 1930 about 4 km south of haredal. it was originally known as djevlekløft and later hermansbu. hareelv 70ø-132 (70°42.3´n 22°44.1´w; map 4). river in eastern jameson land, on the west side of hurry inlet. named by alfred rosenkrantz during lauge koch’s 1926–27 expeditions in the form hare elv, after the arctic hares (fig. 45). harefjeld 80ø-84 (80°16.0´n 20°41.0´w; map 4). mountain in kronprins christian land, on the ne side of vandredalen. named during lauge koch’s 1952–53 expedition by erdhardt fränkl, after the arctic hares. harefjeldet 76ø-67 (76°46.3´n 18°46.8´w). hill 177 m high in southern germania land, on the west side of danmarkshavn. so named by the 1906–08 danmark-ekspeditionen because flocks of hares were often to be seen here. the hill apparently corresponds to the position of karl koldewey’s original kap bismarck, but this name was moved by the 1906–08 danmark-ekspeditionen to the long, low tongue of land se of danmarkshavn. (harefjeld, hare fjæld, hare fjeld, hare hill). harefjord 70ø-1 (70°55.0´n 28°00.0´w; maps 3, 4). fjord running west from the north end of rødefjord, named by carl ryder’s 1891–92 expedition. it was named after the arctic hares. the green landic form ukattit kangersuat has been recorded (tuborg & sandell 1999). haregletscher 72ø-341 (72°27.4´n 22°08.5´w). glacier on se traill ø, draining ellemandsbjerge. named by h.p. heres during lauge koch’s 1956–58 expeditions for the presence of hares. haregletscher 73ø (73°20.7´n 27°19.7´w). name used by the 1972 uni versity of dundee expedition for the glacier in haredalen, ne frænkel land. hareklöft 70ø (70°42.3´n 22°44.1´w). name used by rosenkrantz (1942) for the ravine in which hareelv flows. hareknoldene 74ø-310 (74°06.0´n 21°14.9´w). small hills on the east side of østelv, east of eskimonæs station, clavering ø. the 190 name originated from the wintering party at eskimonæs during the 1931–34 treårsekspeditionen, and was named after the arctic hares. haremarken 73ø-381 (73°38´n 25°13´w). plain in eastern andrée land between morænedal and grejsdalen, named by erdhardt fränkl during lauge koch’s 1948–50 expedition. a total of 34 hares were shot here for food one summer when the expedition ship was delayed by ice and many parties were running short of provisions. hareskindpynten 80ø-83 (80°33.6´n 19°59.5´w). point on the north side of the inner part of ingolf fjord, kronprins christian land, named by the 1938–39 danske hundeslæde-ekspedition. the expedition had a camp at this location, and presumably shot and skinned a hare. hareskåret 76ø (76°46.9´n 18°48.0´w). name reported by fischer (1983) as used by staff at danmarkshavn for the ravine on the nw side of harefjeldet, between hulesøen and stormbugt (hareskåret = harelip). hareø 73ø (73°46.0´n 20°24.0´w). small island in carlshavn on the east coast of hold with hope, probably identical with mågeungen. the name was used as a botanical reference locality by gelting (1934) who visited the island in 1932 during the 1931–34 treårs eks peditionen. named after the hares. harlech fjeld 72ø-487 (72°12.3´n 24°37.3´w; map 5). mountain 1896 m high on the nw side of bersærkerbræ, north stauning alper. first climbed by john hunt’s 1960 expedition, and named harlech after harlech castle, wales. harlech gletscher 72ø-489 (72°11.1´n 24°39.5´w; map 5). glacier on the north side of bersærkerbræ, sw of harlech fjeld, north stauning alper. named harlech glacier by john hunt’s 1960 expedition. harris fjeld 70ø-135 (70°43.5´n 22°42.3´w). hill about 500 m high in eastern jameson land, between primulaelv and hareelv on the west side of hurry inlet. the name was first used in reports by rosenkrantz (1934) in the form harris fjæld, and was given for thomas (tom) maxwell harris [1903–83], who ascended the mountain on 2 september 1926 during lauge koch’s 1926–27 ex peditions and brought back from it the first fossils. harris was a distinguished palaeobotanist who was professor of botany at the university of reading from 1935 to 1968. (harris mountain, mt. harris fjæld.) harry’s hump 72ø (72°15.1´n 24°02.6´w). name used by the 1974 joint biological expedition for two conspicuous small hills on the south side of the valley west of mestersvig airfield. hartz fjeld 70ø-41 (70°42.6´n 25°20.1´w). mountain 669 m high near kap leslie, east milne land. named by alfred rosenkrantz during lauge koch’s 1926–27 expeditions in the form hartz fjæld, to commemorate the botanist of carl ryder’s 1891–92 expedition, nikolaj hartz [1867–1937]. (hartz berg.) hartz vig [kangertivatsiaakajik] 70ø-332 (70°26.8´n 21°48.8´w; map 4). bay in southern liverpool land, ne of kap tobin. the colonisation expedition of 1924–25 had given it the name hartz’s havn, and envisaged it as a possible alternative harbour for ships visiting scoresbysund (e. mikkelsen 1925). it was named after nikolaj hartz, who knew the ice conditions from his participation in carl ryder’s 1891–92 expedition, and had taken great trouble to ease the negotations between the colonisation expedition and the ministry. the bay proved to be often blocked with ice and was never used as a harbour, and the name was later changed to hartz vig. see also hartz fjeld. (hartz havn.) hasdal 72ø-113 (72°40.4´n 25°22.2´w; map 4). valley in lyell land draining east into polhem dal. the name was an adaption by the place name committee of a proposal by eugène wegmann in 1935. wegmann’s original suggestion was thought to be a personal name. hasentinde 72ø (72°01.4´n 24°47.0´w; map 5). summit 2376 m high on the east side of upper storgletscher, central stauning alper. climbed and named by the 2007 smc east greenland expedition. häsi bjerge – see page 200 (in danish ä is treated as æ). haslum øer [traill-iup immikkeertivi] 72ø-57 (72°27.9´n 24°05.5´w; maps 4, 5). group of islands off the sw coast of traill ø. they were named haslums öar by a.g. nathorst’s 1899 expedition, after h.j. haslum [b. 1856] the first mate on the expedition ship antarctic. (haslum island, haslumöyane.) hasserishytten 76ø-207 (76°15.0´n 20°24.5´w). danish hunting hut on the south point of nanok ø, built by nanok in september 1938. it was named after hasseris, a suburb of ålborg, denmark. the hut has also been known as sydlige jægersundhytte. (hasseriis hytten, hasseris hytten.) hastings gletscher 77ø-123 (77°11.8´n 24°37.3´w; map 4). glacier in nw dronning louise land. named by the 1952–54 british north greenland expedition after the hastings aircraft of the royal air force, which air-dropped fuel and equipment to the expedition. one of the aircraft crashed near the ‘northice’ station west of dronning louise land. hastværkshytten 73ø (73°41.3´n 25°06.2´w). name often used for the norwegian hunting hut built by arktisk næringsdrift in 1938 on the north side of morænedal, ne andrée land. it is also known as morænedalshytten. one of the norwegian hunters recorded it as the worst hut he had ever used (p.s. mikkelsen 1994); ‘hastværk’ means a rushed job, implying the hut was poorly built. (villa hastværk.) haugneset 72ø (72°41.2´n 22°06.0´w; fig. 14). small peninsula in extreme se geographical society ø. so named on the nsiu maps of lacmann (1937) after henry georg haug [b. 1907], a nor wegian telegraphist who was stationed at myggbukta in 1934–35 and 1936–37. haugöya 73ø (73°45.2´n 20°29.5´w). island in the delta at the mouth of the river draining into carlshavn, hold with hope. the name appears on the nsiu (1932a) map, and commemorates hen ry georg haug, a norwegian telegraphist. see haugneset. haussman gletscher 74ø-159 (74°02.5´n 22°33.1´w). small glacier in the nørlund alper, northern hudson land, draining north into wordie bugt. first used during lauge koch’s 1929–30 expedifig. 45. arctic hares are common throughout northern east greenland. photo: jakob lautrup. 191 tions in the form haussmann gletscher by backlund (1932). haven 70ø (70°27.0´n 26°15.3´w). area on the west slope of hekla havn, danmark ø, where vegetation is particularly rich. the name is found in one of the reports of carl ryder’s 1891–92 expedition (have = garden). havgrimfjellet 73ø (73°27.6´n 23°26.5´w). mountain 1283 m high on the south side of gauss halvø, corresponding to part of the present smith woodward bjerg. so named on the nsiu (1932a) map after hafgrim (or havgrim), one of the original norse settlers of greenland. havlitsø 70ø-416 (70°29.3´n 27°56.7´w). small lake on sw milne land near rødefjord. named during the 1967–72 ggu scoresby sund expeditions by svend funder after the numerous long-tailed duck (= havlit). havlitsø 76ø-242 (76°49.1´n 19°02.2´w). lake on winge kyst in southern germania land. so named by the 1906–08 danmarkekspeditionen after the long-tailed duck, a common breeding bird in the region. havna 72ø (72°13.7´n 23°45.3´w). norwegian hunting station east of noret on the south side of kong oscar fjord, built by søren richter’s expedition in 1939. named after a small bay below the station known as havna or hamna (= harbour). an earlier hut near the site known as solstrand was moved in 1955. the station was manned from 1939 to 1940 and 1946 to 1951, and subsequently often used as a weekend hut by personnel from mestersvig airfield. it has also been known as trønderheim. (havnahytte, hamna hytte, hamnahytten, hamna hut, hauna.) havnevig 70ø (70°43.7´n 22°38.1´w). bay on the south side of constable pynt [nerlerit inaat], where ships anchor to discharge cargo for the constable pynt airfield. the name is used in the ‘grøn lands havnelods’ (kms 1990). havnhytta – see elveidet. hawkins vandfald 76ø (76°01.4´n 20°09.6´w). waterfall 15–20 m high on the north side of bessel fjord, nw of trums ø, where ships can readily take on water. the name is used in den grøn landske lods (1968). hawley skær 72ø-327 (72°30.5´n 24°15.3´w). skerry west of haw ley ø in holm bugt, sw traill ø. hawley ø 72ø-326 (72°30.5´n 24°14.8´w). island in holm bugt, sw traill ø. the name was proposed by søkortarkivet following their 1956–57 surveying of the channel through vega sund as an alternative approach for ships en route to mestersvig airfield. haystack 75ø-1 (75°43.7´n 19°23.7´w; maps 2, 4). prominent peninsula 305 m high on the east coast of dronning margrethe ii land, north of roseneathbugt, with a conical profile viewed from north and south. named by douglas clavering in 1823 as haystack or the haystack, because of its characteristic shape. originally thought to be an island, it was shown by karl koldewey’s 1869–70 expedition to be connected to the mainland. it is a conspicuous landmark, despite its modest height, and figures as a geodetic marker in many surveys. the difference in position as measured by the 1869–70 koldewey expedition and the 1906–08 danmarkekspeditionen was said to be one of the factors that led to alfred wegener’s theory of continental drift. wegener took part in the 1906–08 danmark-ekspeditionen. (haystock-insel, cape haystack, kap haystack.) haystack-tangen 75ø (75°44.3´n 19°27.6´w). norwegian hunting hut on the low neck of the haystack peninsula, built by john giæver’s expedition in november 1932 (haystackhytten.) hecate glacier 71ø (71°54.8´n 25°36.1´w; map 5). tributary to spærre gletscher in the stauning alper, named by james clarkson’s 1961 expedition after the greek goddess. berchtesgadener gletscher is used for the same glacier in german mountaineering reports. hecla 71ø (71°56.7´n 25°08.1´w; map 5). peak about 2400 m high at the head of cantabræ, stauning alper. so named by the 1998 scottish mountaineering club expedition. heden 70ø-100 (70°48.0´n 24°04.0´w). low-lying coastal stretch of western jameson land. named during the 1931–34 treårseks peditionen by laurits bruhn for its appearance (hede = moor, heath). heden 74ø (74°28.4´n 20°32.9´w). area ne of zackenberg forsk nings station. the name has been used by visiting scientists. heens fjell 71ø (71°54.8´n 25°13.6´w; map 5). mountain about 2530 m high on the north side of roslin gletscher, between ravnas bre and baltos bre. the southern of three summits was climbed by the 1996 norwegian stauning alper expedition, and so named after arner randers heen [1905–1991] of åndalsnes, one of the norwegian climbers who made the first ascent of norsketinden in 1954. heeringhus 76ø (76°44.9´n 18°26.2´w). hut built by danmarkshavn weather station personnel east of the station in the autumn of 1949, on a small island south of øksebladet. it was named after e. heering-hansen, chief mechanic at the station (thomsen 1966). it is also known as øksebladet. heidelbeerberge 73ø (c. 73°28´n 25°22´w). locality near eleonore bugt where the germania ran aground during koldewey’s 1869– 70 expedition. opportunity was taken to carry out scientific investigations, and the name was used in reports (e.g. müller 1974) because of finds of edible berries (j. løve, personal communication 2010). heidrunvatnet 74ø (74°20.2´n 21°25.5´w). lake on theodolit pla teau on west clavering ø. so named on nsiu maps of lacmann (1937), after the goat of old nordic mythology which stood on the roof of the valhal, eating the leaves of a tree. heimdalbreen 74ø (74°18.4´n 21°05.9´w). glacier on central clav ering ø draining east into skillegletscher. so named on nsiu maps of lacmann (1937), after heimdal of old nordic mythology, who was born of nine maidens. heimen 71ø (71°37.8´n 22°59.8´w). norwegian hunting hut on the west coast of wegener halvø, in the inner part of fleming fjord. it was built in the autumn of 1932 by helge ingstad and normann andersen, and was their main depot and hunting station (heimen = the home). it has also been known as ingstadheimen. heimland havn 74ø-83 (74°33.5´n 19°09.5´w). bay on the west side of sabine ø, named by j.m. wordie’s 1926 expedition after the expedition ship heimland, which used the bay as an anchorage. (heimlands havn.) heindalen 74ø (74°15.7´n 20°25.7´w). small valley on east clavering ø. used on the nsiu maps of lacmann (1937), and named after heinir, a poetic old norwegian expression for the inhabitants of hedmarksfylke. heinkel gletscher 75ø-46 (75°10.0´n 22°55.0´w; map 4). glacier at the head of grandjean fjord. mapped and named by lauge koch during flights in 1932 on the 1931–34 treårsekspeditionen. it was named after the three-seater heinkel seaplanes used for surveying flights in 1932 and 1933. (heinkels gletscher.) heintz bjerg 73ø-725 (73°21.0´n 24°38.0´w). mountain about 1500 m high in gunnar andersson land, north ymer ø. named by peter friend following his 1968–70 expeditions after anatal heintz [d. 1975], a vertebrate palaeontologist who had worked in both spitsbergen and greenland, and was director of the palæon to logisk museum, oslo. heivatnet 72ø (72°55.3´n 22°20.1´w). lake on east geographical society ø. used only on nsiu maps (lacmann 1937), and named for its elevated position (heivatnet = high water). hekla havn 70ø-65 (70°26.9´n 26°14.7´w; map 4). shallow, sheltered bay on the south side of danmark ø. named by carl ryder’s 1891–92 expedition after the expedition ship hekla, as the bay was its first place of anchorage since leaving copenhagen and subsequently became the winter harbour. the hekla, registered in tønsberg, was a 240-ton barque-rigged auxiliary steam whaler, built in 1872. later it was purchased by the 1902–04 scottish 192 national antarctic expedition and renamed scotia. during the 1914–18 war the scotia was lost by fire in the scilly isles while operating as an ice-patrol vessel. cairns at the mouth of hekla havn were built by ryder’s expedition, and by members of j.b. charcot’s expeditions in the 1930s. the harbour was apparently known during the expedition under the name kehlers havn. (hekla harbour.) hekla sund 80ø-3 (80°12.5´n 19°00.0´w; maps 1, 4; fig. 24). sound running north and west of lynn ø, south of holm land. so named by the 1906–08 danmark-ekspeditionen after the ship hekla. see hekla havn. (hecla sund.) heklalandet 70ø (70°30´n 26°15´w). name occasionally used by ragnvald knudsen in his diaries of carl ryder’s 1891–92 expedition for the present danmark ø in the inner part of scoresby sund (giæver 1937). the hekla was the expedition ship. see also hekla havn. heklas hvalrossnæss 74ø (74°16.8´n 20°09.0´w). name used for kap berghaus, sw wollaston forland, by ragnvald knudsen during the first visit by norwegian sealers to east greenland in 1889. so named because the crew of the hekla shot 100 walrus on the beach here in half an hour on 16 july (knudsen 1890; solberg 1929; giæver 1937). see also hekla havn. a few walrus still come ashore regularly on nearby sandøen. (heklas hvalrosnæs.) heksefjeldet 70ø-173 (70°39.1´n 22°10.4´w). mountain about 800 m high in southern liverpool land between hans gletscher and grete gletscher. named during the 1931–34 treårsekspeditionen by laurits bruhn with the two glaciers after the characters in the grimm brothers’ fairy tale ‘hänsel und gretel’ (hans og grete in danish; hansel and gretel in english; heks = witch). helgegletscher 70ø-78 (70°18.0´n 25°02.4´w). glacier on the south side of scoresby sund, on the east side of vikingebugt, entering the sea just west of helgenæs. so named by laurits bruhn during the 1931–34 treårsekspeditionen after helgenæs. helgenæs 70ø-80 (70°21.8´n 25°02.0´w; map 4). peninsula on the south side of scoresby sund, east of vikingebugt. named by laurits bruhn during the 1931–34 treårsekspeditionen after the peninsula of the same name east of aarhus in jylland, denmark. helgoland 76ø-134 (76°26.9´n 26°20.2´w; map 4; fig. 21). nuna tak 2125 m high in sw dronning louise land. named by j.p. koch’s 1912–13 expedition after the island of the same name off the nw coast of germany. it was climbed by the lancaster uni versity expedition in may 2000. helispids 71ø-411 (71°59.3´n 23°03.0´w). mountain 838 m high south of antarctic havn, ne scoresby land. so named by kathe rina perch-nielsen during the 1967–72 ggu scoresby sund expeditions because it was ‘climbed’ by helicopter. hellandfjellet 74ø (74°00.8´n 22°45.4´w). mountain ridge in the nørlund alper, north hudson land. so named on nsiu maps of lacmann (1937), after amund helland [1846–1915], a norwegian geologist who visited west greenland in 1875. hellas strømhvirler 70ø-382 (70°14.5´n 28°59.7´w). whirlpool in the turbulent river between gnejssø and kaskadesø, western gåse land. named during the 1958 lauge koch expedition by eduard wenk, for an incident in which his greek assistant, j. papage orgakis, was nearly drowned. derived from hellas, the greek name for greece. helledalen 76ø-218 (76°48.4´n 21°26.0´w). valley crossing daniel bruun land from hellefjord to port arthur. the name was sug gested by the place name committee to replace a proposal by the 1938–39 mørkefjord expedition. hellefjeld 72ø-263 (72°13.9´n 25°17.1´w; map 5). mountain 1947 m high in the northern stauning alper, north of the front of vikingebræ. it was first climbed by the dansk–norsk grøn landsekspedition in august 1954, and named at the suggestion of john haller, possibly after the german word for light (= hell) recording the light colour of the rocks. hellefjord 76ø-23 (76°51.3´n 21°09.4´w; map 4). wide and open fjord extending westwards into daniel bruun land. named by the 1906–08 danmark-ekspeditionen, possibly for the contrast with the narrow enclosed mørkefjord, and perhaps derived from the german word for light (= hell). the expedition included the ger man scientist alfred wegener. (helle fjord, helle fiord, hellu fjorður.) hellefjordhytten 76ø-195 (76°56.7´n 21°21.6´w). danish hunting hut on the north side of inner hellefjord, western dove bugt, built by nanok in september 1933. it is now a ruin (1989). helmspitzen 71ø (71°58.0´n 25°14.7´w; map 5). mountain about 2400 m high on the sw side of sefström gletscher, stauning alper. climbed by the 1964 aac zürich expedition, and named for its helmet-like shape. helvedespas 72ø-253 (72°10.2´n 24°55.7´w). pass between vikingebræ and skjoldungebræ, north stauning alper. named during lauge koch’s 1951 expedition by erdhardt fränkl, the first approach being made by peter braun and fritz schwarzenbach from the skjoldungebræ side in an attempt to reach dansketinden. they failed to climb dansketinden on this occasion because they could not reach ‘the damned pass’. ‘helvedes’ is a danish expletive (helvede = hell). hendil valley 73ø (73°30.0´n 27°22.1´w). name used by noel e. odell during louise boyd’s 1933 expedition for a valley just north of kap hendil, louise boyd land (odell 1939). henius nunatak 77ø-51 (77°08.7´n 25°03.5´w; fig. 21). nunatak in nw dronning louise land, so named by the 1909–12 alabama expedition after erik semmy henius [1863–1926], a member of the expedition committee. henius was a danish consul and businessman noted for his interests in arctic research. hennigryggen 70ø-43 (70°42.0´n 25°19.5´w). ridge between kostenbaderbjerg and slottet, nw of kap leslie, east milne land. named by hermann aldinger during the 1931–34 treårseks peditionen as hennigrücken or hennig berg, probably after edwin hennig [b. 1882], a german palaeontologist and stratigrapher noted for his work in africa. henning dal 76ø-147 (76°48.2´n 21°49.3´w; map 4). valley in west daniel bruun land, named by j.p. koch’s 1912–13 expedition as hennings dal. it was probably named after henning bistrup [1879–1948], one of the founders of østgrønlandske fangst kompagni, a member of the 1906–08 danmark-ekspeditionen and captain of the teddy in 1923. (henningsdalur.) henning valley 74ø (74°12.6´n 20°15.4´w). name occasionally used by danish hunters for the valley on east clavering ø in which henningelv flows. see henningelv. henningelv 74ø-109 (74°12.6´n 20°15.4´w). stream on east clavering ø flowing north into young sund. the name first appeared in the form henningselv on a sketch map in gustav tho strup’s 1921 logbook, but was there applied to a river flowing east to enter the sea just south of kap arnakke. the name was subsequently commonly used by danish hunters for the present river, and probably commemorates henning bistrup. see also henning dal. (henning river.) henningelvhytten 74ø (74°13.4´n 20°14.0´w). danish hunting hut on the west side of the mouth of henningelv in east clavering ø. built by nanok in july 1930, and renovated by sirius in 1993. henrik kröyer holme 80ø-12 (80°38.3´n 13°43.2´w; maps 1, 4). group of three low islands se of amdrup land, named by the 1906–08 danmark-ekspeditionen as henrik kröyers holme, after henrik nikolaj krøyer [1799–1870]. krøyer was a danish zoologist who travelled widely, including voyages to south america and spitsbergen, and was noted particularly for his ‘danmarks fiske’, published from 1838 to 1853. the islands were first visited by gustav thostrup and alfred wegener in april 1907. the islands are a notable breeding area for birds; 300 pairs of ivory gulls, more than 100 pairs each of arctic tern and common eider, 50 pairs of 193 sabine gull and smaller numbers of other birds were noted in 1993. an automatic weather station was erected on one of the islands in july 1984. (henrik kröyers islet.) henrik møller dal 71ø-169 (71°52.7´n 22°57.6´w). valley north of the mouth of ørsted dal. the name was one of a group of names given by the place name committee in 1939. it commemorates the danish civil servant henrik møller, head of the customs administration, who promoted david danel’s three voyages to greenland from 1652–54. henry bjerg 69ø-26 (69°34.0´n 23°44.0´w). name used by böggild (1905) in his mineralogical description of g.c. amdrup’s 1898– 1900 rock collections, which was used in the form mount henry or henry mountain. the name was probably intended for the southern large peninsula of henry land, on the northern blosseville kyst. see also henry land. (henry bjærg.) henry glacier 69ø (69°38.0´n 24°04.0´w). name used by böggild (1905) in his mineralogical report of g.c. amdrup’s 1898–1900 expedition for the glacier sw of henry land now known as bar tholin bræ. henry land 69ø-7 (69°40.0´n 23°54.0´w; map 3). land area between rømer fjord and bartholin bræ on the northern blosse ville kyst. william scoresby jr. named henry island in 1822 after dr. william henry [1774–1836], a prominent chemist, who had studied medicine at edinburgh university at the same time as scoresby. scoresby’s island was later discovered to be a peninsula, the name henry land being first used by hartz (1902) and koch (1902). (henry ö, henry peninsula, henry halvø.) herdal 73ø (c. 73°33´n 22°56´w). name used in a report by nsiu (1932c) for a prospective hut on the south side of moskusoksefjord at the mouth of västidal. material for the hut was deposited here in 1930 by arktisk næringsdrift, but the hut was never built (p.s. mikkelsen 2008); the material was used later for petrahytten. the hut name was given for eilif herdal [b. 1897], a norwegian hunter who took part in a hunting expedition to east greenland from 1929 to 1932, and led his own expedition in 1933–35. herdalfjellet 72ø (72°59.3´n 23°49.7´w). mountain 960 m high on western geographical society ø. used only on nsiu maps (lac mann 1937), and named after eilif herdal. see also herdal. herjaelv 73ø-197 (73°44.3´n 21°39.5´w). river on the east side of loch fyne, named on the nsiu (1932a) map in the form herja. derived from the norwegian word for something violent or strong. herjahytten 73ø (73°44.0´n 21°40.5´w). danish hunting hut in hold with hope on the south side of herjaelv, about 4 km inland. built by nanok in august 1938. (herjadalshytten.) herman andresenfjellet 74ø (74°09.9´n 20°52.7´w). mountain 1330 m high on south clavering ø, equivalent to the present pladen. so named on nsiu maps of lacmann (1937), after herman andresen [b. 1901], a norwegian hunter who spent several years in east greenland between 1927 and 1940. from 1947 to 1959 andresen organised a succession of norwegian hunting expeditions to east greenland. he was regarded by norwegian hunters as their greatest story teller. hermann von barth tinde 71ø (71°54.2´n 25°41.1´w; map 5). moun tain 2681 m high on the west side of spærregletscher in the western stauning alper, at the head of pollux glacier. named after the noted german mountaineer hermann von barth [1845–1876], born in berchtesgaden. first climbed by the 1967 berchtesgadener expedition. hermansbu 74ø (74°18.1´n 19°18.9´w). norwegian hunting hut on the east coast of wollaston forland about 4 km south of haredalen. built in july 1930 by the møre expedition, and named after the norwegian hunter herman andresen, who helped build it. see herman andresenfjellet. it was originally called djevlekløft, and has also been known as haredalshytten. hermelintop 70ø-396 (70°26.3´n 27°56.9´w; map 4). summit 1172 m high on sw milne land. so named during the 1963 geodætisk institut expedition because ermine (= hermelin) were observed here during surveying (fig. 46). hermes 71ø (71°37.1´n 25°10.3´w; map 5). mountain about 2100 m high on the south side of mercurius gletscher, south stauning alper. first climbed by james clarkson’s 1961 expedition, and named after the greek god, son of zeus and maia. hermitage 71ø (71°47.1´n 25°01.3´w; map 5). mountain about 2200 m high at the head of mars gletscher, south stauning alper. first climbed by james clarkson’s 1961 expedition, and probably named after hermitage castle, a remote 13th castle on the scottish borders. hermodbreen 74ø (74°17.3´n 21°02.5´w). glacier on central clav ering ø, draining east into skillegletscher. so named on nsiu maps of lacmann (1937), after hermod, son of odin in old nordic mythology. heron hump 70ø (70°46.5´n 22°01.1´w). minor summit 788 m high in liverpool land, on the north side of bjerring pedersen glet scher. it was climbed and named by the 2002 loughborough grammar school expedition. herschell bjerg 74ø-8 (74°16.1´n 19°42.3´w; map 4). mountain fig. 46. ermine (hermelin) are only seen in large numbers in so-called lemming-years, when the abundance of lemmings boosts the breeding popu lation of ermine, foxes and snowy owls. 194 682 m high in south wollaston forland. it was named cape herschel by william scoresby jr. in 1822 after john frederick william herschel [1792–1871], baronet, physicist and astron omer, noted for his survey of the skies in the southern hemisphere. like many of scoresby’s capes it was observed from a great distance and the name was later transferred to the mountain he had probably seen. scoresby misspelt the name as cape herschell on the maps in both english (1823) and german (1825) editions of his narrative, and it is this spelling that has been used on virtually all maps to the present. it was commonly referred to as kapp herschel in the 1930s in association with the norwegian hunting station her schell hus at its foot. danish hunters have used etagefjældet for the same feature. (mt hershell, herschelfjellet.) herschellhus 74ø-243 (74°14.6´n 19°41.1´w). norwegian hunting station south of herschell bjerg, southern wollaston forland. originally built by the hird expedition in 1927, it was improved and enlarged in 1929, 1930 and 1952. the name appears on the nsiu (1932a) map as herschelhus, and is often referred to in hunting accounts as kapp herschel. it was manned almost continuously in the periods 1927–41 and 1946–57. herthabjerg 72ø-473 (72°10.8´n 26°54.8´w). mountain 1910 m high on the north side of the mouth of herthadal, nathorst land. named by hans zweifel during lauge koch’s 1954–55 expeditions. see also herthadal. herthadal 72ø-435 (72°10.3´n 27°04.9´w; map 4). valley west of violingetscher, nathorst land. named during the 1931–34 treårs ekspeditionen by ove simonsen, and given for a danish locality of the same name near roskilde, sjælland. hertugen 77ø-125 (77°08.3´n 24°54.1´w; map 4; fig. 21). high, dark peak in nw dronning louise land, ne of prinsessen. named by the 1952–54 british north greenland expedition after prince philip, duke of edinburgh [b. 1921], husband of the expedition patron queen elizabeth ii. prinsessen and hertugen (= the duke) are the two highest peaks in northern dronning louise land. hertugen was climbed by mike banks and malcolm slesser in 1953, who reported finding the bones of a wolf near the summit. hertugen af orléans land 77ø-97a 78ø-1 (77°34´n 22°00´w to 78°45´n 21°15´w; maps 1, 2, 4). land area on the west side of jøkelbugten. named by the 1905 the duke of orléans expedition as terre du duc d’orléans, although originally used in a wider sense to include the offshore islands (fig. 9). the duke of orléans had wanted to call the region terre de france, but the danish administration vetoed this suggestion. the 1906–08 danmark-ekspedi tionen moved the name to the present location after consultation with adrien de gerlache (koch 1916 p. 376), where it includes sønderland, søndre mellemland, mellemland and nørreland. louis-philippe-robert, duke of orléans [1869– 1926] had served with the british army in the west indies, and on the death of his father became pretender to the french throne. he made a number of voyages to the arctic. hesleyside 73ø (73°31.8´n 25°55.0´w). summit 2099 m high on the south side of grejsdalen, andrée land. climbed by the 2007 army boreal zenith expedition. hessbreen 74ø (74°01.8´n 22°26.1´w). glacier in the nørlund alper, north hudson land, equivalent to a. schmidt gletscher. used on the nsiu maps of lacmann (1937), it was named after hans hess [1864–1940], a german glaciologist. hesselbergfjellet 72ø (72°56.0´n 24°26.2´w). mountain 1677 m high on west geographical society ø, corresponding to part of the present svedenborg bjerg. used only on nsiu maps (lacmann 1937), it was named after hans theodor hesselberg [1885–1966], a director of the norwegian meteorological institute from 1915. hesselberg was responsible for the norwegian weather stations in greenland. hesteelv 70ø-103 (70°28.1´n 22°58.5´w; map 4). river in south jameson land, named by hermann aldinger during the 1931–34 treårsekspeditionen as horse river. it commemorates the use of icelandic ponies (heste = horse) during the geological exploration of jameson land. hestefoden 76ø-16 (76°24.8´n 20°19.8´w). the curved horse-shoe shaped northern part of djævleøen, dove bugt, so named by the 1906–08 danmark-ekspeditionen. the devil (= djævle) is alleged to have had hooved feet (= hestefod). hestehale sø 74ø (74°29.3′n 20°36.5′w). small lake in the sw part of morænebakkerne, north of zackenberg forskningsstation. the name has been used as a reference locality in ornithological reports by visiting scientists. hestepas 72ø-223 (72°10.0´n 23°47.3´w; map 5). low col beside myggesø, west of the mouth of mesters vig. so named by prospecting teams associated with lauge koch’s 1948–49 expeditions, because it lies on the pony route between the airfield at mestersvig and expeditionshus. (hestepasset.) hesteskoen 71ø-91 (71°38.8´n 22°20.9´w). mountain on canning land, named during the 1931–34 treårsekspeditionen by arne noe-nygaard as hesteskofjeld because of its shape (hestesko = horse shoe). hesteskoen 72ø-182 (72°13.1´n 24°11.1´w; map 5). mountain in northern scoresby land, east of skeldal. named by prospecting teams associated with lauge koch’s 1948–49 expeditions for its horse-shoe like shape. heywood bjerge 70ø-220 (70°41´n 21°44´w). mountain massif between kolding fjord and lille fjord on the coast of liverpool land. named originally as heywood island by william scoresby jr. in 1822 in compliment to a mr. b.a. heywood. (heywood insel.) hidden valley 73ø (73°21.7´n 25°11.1´w). valley between choko lade bjerg and rosinante, west ymer ø, the present rosinante pas. the name was given by arthur b. cleaves and ernest f. fox in the course of geological work during john k. howard’s 1933 expedition (cleaves & fox 1935), because the valley was hidden by little chocolate mountain (now rosinante). highgate 72ø (72°04.2´n 24°39.5´w; map 5). mountain 2450 m high at the head of kishmul gletscher, north stauning alper, the present kishmul borg. first climbed by the 1963 imperial college expedition, and named after the north london district of high gate, which originally had a toll gate on top of a hill. highway – see the highway. hildebrandbreen 74ø (74°19.5´n 21°17.7´w). glacier on central clavering ø, draining to the west. so named on the nsiu maps of lacmann (1937) after hildebrand, who features in the german epic poem from c. 1200, the nibelungenlied. hildegard island 71ø (71°16.6´n 21°42.4´w). name used occasionally in reports of the 1931–34 treårsekspeditionen (e.g. kranck 1935) for the present island trekanten, liverpool land. the name was given by helge g. backlund for his wife hildegard dischner, whom he married in 1914. two nearby capes, kap hilding and kap vidar, were named after his sons. hill end pond 72ø (72°14.4´n 23°55.0´w). name used by the 1974 joint biological expedition for a pool near langdyssen at the ne end of mestersvig airfield. himmelberget 74ø (74°13.2´n 20°17.4´w). mountain 241 m high on east clavering ø, on the west side of henningelv. the name is used by lacmann (1937), and may be a reference to himmelbjerget in jylland, denmark. norwegians traditionally make fun of the diminutive height (175 m) of this danish ‘mountain’. himmelbjerg 73ø-359 (73°49.4´n 24°39.3´w; map 4). mountain about 1400 m high in strindberg land, named by hans r. katz during lauge koch’s 1948–49 expeditions (himmel = sky). (himmelbjærg.) himmelpas 73ø-360 (73°49.1´n 24°43.5´w). high pass between rævedal and rodedal, strindberg land. named during lauge koch’s 1948–49 expeditions by hans r. katz. himmelstinde 72ø (72°04.9´n 25°05.4´w; map 5). peak 2492 m high 195 on the west side of upper gullygletscher, northern stauning alper. climbed and named by the 2007 smc east greenland expedition; the name was translated as ‘heavens peak’. himmerland 70ø-254 71ø-130 (71°02´n 21°55´w). peninsula between mariager fjord and storefjord, liverpool land. so named during the 1931–34 treårsekspeditionen by laurits bruhn after the district of the same name in jylland, denmark. himmerland hede 76ø-125 (76°41.0´n 24°00.0´w; map 4). pla teau on the north side of borgjøkelen, dronning louise land. named by j.p. koch’s 1912–13 expedition as himmerlandshede, after the area of the same name in denmark where one of the expedition members, lars larsen, was born (hede = heath). (himmer lands hede, himmerlandsheiði.) hindarfjellet 74ø (74°22.5´n 21°03.4´w). mountain ridge about 1430 m high on north clavering ø, ne of ortlerspids. the name appears on the nsiu maps of lacmann (1937), and was named after a character in the german epic poem from c. 1200, the nibe lungenlied. hindringsgletscher 73ø-417 (73°59.6´n 28°02.8´w). glacier between bernhard studer land and arnold escher land; so named by hans r. katz during lauge koch’s 1951 expedition because it was an obstacle (= hindring) to their progress. hinks land 71ø-64 (71°40.0´n 28°30.0´w; maps 3, 4; fig. 41). land area between daugaard-jensen gletscher and flyverfjord. the name first appeared on the 1932 1:1 million scale geodætisk institut map prepared on the basis of 1932 aerial observations by lauge koch during the 1931–34 treårsekspeditionen. the name was given for arthur robert hinks [1873–1945], a british mathematician and an authority on map projections, and the very influential secretary of the royal geographical society from 1915 to 1945. hird bay 74ø (74°08.7´n 20°33.3´w). open bay on se clavering ø, west of basaltkap. named by lauge koch’s 1929–30 expeditions after the main hunting station (elvsborg ) constructed by the 1927–29 hird expedition on the west side of the bay. see also hirdhavn. (hirds bay.) hird star 71ø (71°48.6´n 24°59.0´w). prominent peak on the south side of roslin gletscher, about 2159 m high. so named and climbed by the 1970 university of cambridge expedition on 15 august 1970, the 3rd ascent. probably named after t.a. hird, a member of the 1968 queen mary college expedition, who was evacuated by helicopter after falling into a glacier stream. the first ascent of the mountain was by karl herligkoffer’s 1966 expedition, which had called it granit spids. the second ascent in 1970 was by a university of dundee party. hirdhavn 74ø-265 (74°03.0´n 20°52.1´w). small bay or harbour on the north side of store finsch, the largest island of the finsch øer. the hird, a 48 foot fishing boat used by the 1927–29 hird expedition, was anchored in the bay for the winter, but was wrecked and sank in a storm on 27 august 1927. norwegian hunters used the form hirdbukta or hirdhamna. hirds fox farm 74ø (74°07.9´n 20°39.9´w). hunting station built in 1927 by the norwegian 1927–29 hird expedition in se clavering ø, also known as elvsborg. the name is encountered in a number of expedition reports, and is a reference to the practice of keeping trapped foxes alive in cages at the station until their fur was in optimal condition. see also hirdhavn. (hirds rævefarm.) hirschbichler spids 71ø (71°56.9´n 25°39.8´w; map 5). mountain on the west side of spærregletscher, western stauning alper, between castor glacier and pollux glacier. named and first climbed by the 1967 berchtesgadener expedition. hisinger gletscher 72ø-401 (72°49.0´n 27°38.0´w; map 4). gla cier at the head of dickson fjord, between suess land and gletscherland. named by a.g. nathorst’s 1899 expedition, probably after wilhelm hisinger [1766–1852], a chemist and mineralogist who with j.j. berzelius had discovered the element cerium. hisinger’s interests were mainly geological, and his collections form the basis of the rijksmuseum mineralogical collections in stockholm. (hisingers glacier.) hjelmbjergene 73ø-101 (73°28.0´n 23°28.1´w; map 4). range of mountain summits on the se coast of gauss halvø. the name was adopted from a suggestion by th. johansen during the 1931–34 treårsekspeditionen, who had likened them to the helmets of a line of roman soldiers. hjelmen 72ø-120 (72°52.2´n 25°59.3´w). snow-capped mountain 2152 m high in southern suess land, west of kap buxtorf. so named by the 1931–32 ella ø wintering party during the 1931–34 treårsekspeditionen for its helmet-like appearance. hjelmen 74ø (74°16.5´n 21°49.7´w). mountain on west clavering ø, equivalent to the present dunken. used only on nsiu maps (lacmann 1937), and named for its helmet-like shape. hjelmen 76ø-143 (76°34.5´n 25°07.6´w; map 4). mountain in sw dronning louise land, named by j.p. koch’s 1912–13 expedition for its rounded form resembling a helmet. hjerte sø 74ø (74°30.5′n 20°37.8′w). small heart-shaped lake in the area known as morænebakkerne, north of zackenberg forsknings station. the name is used as a reference locality by scientists studying lake ecosystems. hjertet 74ø-133 (74°15.2´n 20°58.5´w). mountain about 1400 m high on central clavering ø. the name was first used by mittel holzer (1941), and is a reference to the shape, or possibly the central placing of the mountain (hjertet = the heart). hjort lake 76ø (76°26.0´n 18°45.7´w). lake on store koldewey where samples were taken for radiocarbon age determinations and phytoplankton studies (cremer et al. 2005, 2008). hjælmen 76ø-223 (76°57.1´n 20°47.2´w). mountain west of pyra midedalen on the north side of mørkefjord, daniel bruun land. named by the 1938–39 mørkefjord expedition, probably for the helmet-like shape. ‘hjælm’ is an old danish spelling of ‘hjelm’ (= helmet). hjørnebjerg 72ø (72°10.8´n 26°54.8´w). name used by zweifel (1958), apparently for the mountain north of hjørnesø, nathorst land, of which the peak is known as herthabjerg. hjørnebjerget 74ø-334 (74°02.5´n 23°43.5´w). mountain 1137 m high at the bend of krumme langsø (hjørne = corner, bend). named during lauge koch’s 1936–38 expeditions by heinrich bütler. (hjørneberg.) hjørnedal 70ø-16 (70°19.0´n 28°15.6´w). valley in gåseland draining into the sea where fønfjord meets rødefjord at a right angle. named in this form by carl ryder’s 1891–92 expedition. hjørnefjeld 77ø (77°04.4´n 20°28.0´w). name given by the 1938–39 mørkefjord expedition to the southernmost part of valdemars muren, nw of trekroner, western germania land. it may be identical with the summit above depotkulle. hjørnefjeldet 71ø-57 (71°12.4´n 22°49.3´w). mountain c. 800 m high in eastern jameson land with a curved summit ridge. named during lauge koch’s 1926–27 expeditions by alfred rosenkrantz and tom harris as mt hjørnefjæld. it is misplaced about 10 km farther north on some editions of the geodætisk institut 1:250 000 scale map sheet (71 ø.1). hjørnegletscher 80ø-49 (80°39.7´n 19°26.9´w; map 4). glacier on the north side of inner ingolf fjord, where the fjord makes a right-angled bend. named by eigil nielsen during the 1938–39 mørkefjord expedition. hjørnemoræne 71ø-372 (71°18.7´n 24°53.3´w; map 5). moraine ridge east of sydkap, at the corner between the mouth of schuchert dal and nordvestfjord. named by the 1962 oxford university ex pedition. hjørnepunktet 73ø-50g (73°59.0´n 21°24.0´w). point between wordie kløft and blåelv, nw hold with hope. so named by eigil nielsen as hjörnepunkt during the 1931–34 treårsekspeditionen because of its location at a corner overlooking blåelv. 196 hjørnespids 72ø-323 (72°07.7´n 24°55.7´w; map 5). mountain 2650 m high between the heads of gully gletscher and bersærker bræ, north of majorpasset, north stauning alper. named by john haller in 1957, it is sometimes confused with the mountain pyra midefjeld to the north (bennet 1972). first climbed by john hunt’s 1960 expedition, and subsequently by the 1968 queen mary college expedition. (eckspitze.) hjørnesø 72ø-434 (72°08.1´n 26°51.1´w; map 4). lake between jomfrudal and violingletscher, nathorst land. so named by ove simonsen during the 1931–34 treårsekspeditionen because of its position at a corner of the glacier. hjørnet 72ø-215 (72°07.1´n 24°02.4´w; map 5). mountain north of the mouth of nedre funddal, northern scoresby land. named by prospecting teams associated with lauge koch’s 1948–49 expeditions. hjørnet 73ø-671 (73°39.6´n 26°57.3´w). mountain c. 2000 m high in west andrée land. so named during lauge koch’s 1950 expedition by john haller, because it was in an outlying corner (= hjørne) of the region he mapped. hobbs land 73ø-594 (74°03.0´n 29°00.0´w; map 4). area of nunataks at the west extremity of adolf hoel gletscher. mapped and named by lauge koch during flights in 1932 on the 1931–34 treårsekspeditionen. the name commemorates william herbert hobbs [1864–1952], an american geologist who was professor at the university of wisconsin from 1889 to 1905 and at the uni ver sity of michigan from 1905 to 1934. he led several university of michigan expeditions to west greenland in the 1920s. on modern maps the area is just north of 74°n latitude. hochstetter 75ø (75°08.5´n 19°44.9´w). name commonly used by danish hunters in the 1930s for the danish hunting station in southernmost hochstetter forland, officially known as nanok. (hochstetter station.) hochstetter forland 75ø-31 (75°25.0´n 19°48.0´w; maps 2, 4). low-lying land area ne of ardencaple fjord, limited to the west by the barth bjerge. named by karl koldewey’s 1869–70 expedition as hochstetter vorland, after ferdinand ritter von hochstetter [1829–1884]. an austrian geologist, he was professor in mineralogy in vienna and had coordinated the geological chapter of kolde wey’s expedition narrative. (hochstetter’s promontory, hochstetters forland.) hochstetterbugten 74ø-314 75ø-31a (74°54.0´n 19°00.0´w; maps 2, 4). broad bay between hochstetter forland and shannon to the north, and wollaston forland and the pendulum øer to the south. the name is said to have been in use from 1929 by danish hunters, and first appeared on the maps of the 1932 gefion expedition. (hochstetters bugt.) hochwacht 73ø-326 (73°46´n 23°10´w). mountain 1605 m high in hudson land south of ritomsø. named during lauge koch’s 1936–38 expeditions by heinrich bütler, after the small prominent hills of switzerland used as watchtowers (= hochwacht). hodal 70ø-152 (70°54.8´n 22°24.9´w). valley in liverpool land on the east side of hurry inlet. so named by laurits bruhn during the 1931–34 treårsekspeditionen, possibly after the small town of ho near esbjerg, denmark. hodbreen 74ø (74°22.1´n 21°01.0´w). glacier on clavering ø draining east to the front of skillegletscher. the name is used on the maps of lacmann (1937), and derives from old nordic mythology. høder, the blind son of odin, was lured into shooting his brother balder with an arrow made of misteltoe, the only thing which could hurt him. hoelhuset – see holstad. hoelsbo 73ø-265 (73°42.2´n 23°26.3´w). norwegian hunting station on the north side of moskusoksefjord, 5 km sw of genvejsdal, built by john giæver and otto johnsen in august 1930 for arktisk næringsdrift. named hoelsbu after adolf hoel [1879–1964], a norwegian geologist, and the driving force behind norwegian scientific activities in east greenland. from 1909 onwards, hoel took part in about 30 norwegian government-sponsored expeditions to the arctic. he was director of nsiu (subsequently norsk polar institutt) from 1928 to 1945, and was actively concerned with norwegian acquisition of polar territories, svalbard in 1928 and dronning maud land in the antarctic in 1938. hoelsbo was manned almost continuously in the periods 1930–42 and 1946–59. it has been maintained by sirius, and was still in good condition in 1988, although the floor slopes gently towards the fjord. (hoelsby.) hofgaardvatna 72ø (72°43.6´n 22°29.3´w; fig. 14). lake on se geographical society ø, the present basaltsø. used on the nsiu maps of lacmann (1937), the name was given for knut hofgaard [b. 1903], a norwegian hunter who wintered in east greenland from 1932 to 1933. hohe kugel 74ø-30 (74°41.0´n 20°53.0´w; map 4). mountain 1337 m high on the north side of lindeman fjord, so named by karl kolde wey’s 1869–70 expedition probably after the mountain of the same name in austria. (mt hohe kugel, store kugle.) hohgant 74ø-354 (74°38.4´n 20°08.2´w). mountain 658 m high in northern wollaston forland. so named during lauge koch’s 1936– 38 expeditions by wolf maync and andreas vischer (maync 1947) because in shape and geology it resembles the mountain ridge of the same name in the berner oberland, switzerland. holberg elv 72ø-220 (72°07.2´n 23°55´w; map 5). river draining from holbergpasset eastwards to mesters vig. named by prospecting teams associated with lauge koch’s 1948–49 expeditions after ludvig holberg [1684–1754], a danish historian and writer, who wrote 25 plays for the theatre. on detailed 1:15 000 scale topographic maps it is also referred to by the designation 2v. holbergpasset 72ø-212 (72°09.0´n 23°58.5´w). pass across blyryg gen at the head of store blydal, north scoresby land. named by prospecting teams associated with lauge koch’s 1948–49 expeditions. see also holberg elv. (holberg pass.) holbergs bjærg 73ø (73°26.6´n 22°04.1´w). name used on the maps of maync (1942) for the point 818 m high on the present bonney plateau, giesecke bjerge. it was originally suggested by the place name committee, and approved in 1938, but was later abandoned. see also holberg elv. (holbergs fjeld.) hold with hope 73ø-2 74ø-80a (73°45.0´n 21°00.0´w; maps 2–4). land area between foster bugt and gael hamke bugt, bounded to the west by loch fyne. the name is the oldest place name in east greenland north of 69°n to have survived, and derives from henry hudson’s 1607 voyage in the hopewell of hull, a muscovy company whaler active in the 17th century. hudson described it as a “mayne high land”, a “good land, and worth the seeing” (pur chas 1906 pp. 297–298). the name appears on a dutch map by j. hondius dated 1618 as holde with hope, and has been variously applied to smaller parts of the present area, or to include also gauss halvø and hudson land. it has also been appended to the present kap broer ruys, which appears on a number of maps as cape hold with hope. (hold with hope, hold-with-hope, hold-with-hope land.) holger danske briller [imeq] 71ø-60 (71°25.5´n 25°08.0´w; maps 4, 5). name given to two adjacent large lakes at the south extremity of the stauning alper, north of sydkap. the name first appeared on the 1932 1:1 million scale geodætisk institut map prepared on the basis of 1932 aerial observations by lauge koch during the 1931–34 treårsekspeditionen. the lakes resemble in plan a pair of giant spectacles (= briller), whose size suggests they might be the property of the danish legendary sleeping giant holger danske, said to awaken whenever denmark is in peril. it has been suggested that the name was given as a symbol of protest against norwegian claims to sovereignty over east greenland. holger danske tinde 74ø-140 (74°27.1´n 24°33.6´w). isolated snow covered pyramid 2000 m high in bartholin land. named by lauge koch’s 1929–30 expeditions in the form holger danske peak, 197 and thought originally to be 3000 m high (seidenfaden 1931). see also holger danske briller. holland ø 73ø-11 (73°36.0´n 20°21.0´w; map 4). small island 89 m high off hold with hope. william scoresby jr. had named a feature in this vicinity as cape holland, in 1822. henry holland [1788– 1873] was a physician who graduated from edinburgh university in 1811 where he probably met scoresby. scoresby’s ‘cape’ was seen at a great distance, and may have been a mountain on hold with hope, possibly that between tværelv and orvaelv. the name was transferred to an island by karl koldewey’s 1869–70 expedition. (holland insel, dutch island, cape holland, holland øya). holloway bugt 70ø-244 (70°54.0´n 21°41.6´w; map 4). bay south of kap greg on the coast of liverpool land. named holloway bay by william scoresby jr. in 1822 after his friend, the revd richard holloway, a preacher of evangelical and calvinistic sentiments. scoresby had named his second son (born 1818) frederick richard holloway scoresby. (halloway bugt.) holm bjerg 80ø-18 (80°06.6´n 21°01.5´w; map 4; fig. 24). moun tain about 1430 m high in kronprins christian land, east of centrumsø. named by the 1909–12 alabama expedition as holm’s nunatakker after gustav frederick holm [1849–1940], danish naval officer and polar explorer. holm took part in several expeditions to greenland, notably as leader of the 1883–85 expedition to se greenland which discovered the greenland inuit community at ammassalik. he was a member of the alabama expedition com mittee. in may 1913 ejnar mikkelsen married gustav holm’s daughter, naja marie heiberg holm (j. løve, personal communication 2009). the identification of this particular mountain as the ‘nunatak’ seen by the alabama expedition is somewhat speculative, but the place name committee decided the name should be preserved. holm bugt 72ø-56 (72°30.5´n 24°04.7´w). bay on sw traill ø. named as holms vik by a.g. nathorst’s 1899 expedition, probably after gustaf birger anders holm [1845–1910], a publisher of edu cational books who guaranteed a sum of 2500 swedish kronor in respect of the expedition. a hut was built in the bay by norwegian hunters in 1932 (see holm-vika). (holm bay, holmbukta, holm vika.) holm bugt hytten – see holm-vika. holm land 80ø-5 (80°20.0´n 17°00.0´w; maps 1, 4). land area between ingolf fjord and dijmphna sund, eastern kronprins christian land. it was named by the 1906–08 danmark-ekspedi tionen as holms land, after gustav frederick holm. see also holm bjerg. holm-vika 72ø (72°30.1´n 24°00.3´w). norwegian hunting hut at the head of holm bugt, traill ø, built by helge ingstad’s expedition in july 1932. the hut has been used as a base by the grea karupelv valley project, and was restored by nanok in 2001. it has also been known as karupelv hytten. (holms vig hytten, holmsvik, holmbugta, holm bugt hytten.) holmboe-hytta 73ø (74°27.9´n 20°39.1´w). hut in innermost dusén fjord built for salmon fishing in august 1932 on behalf of f.a.w. holmboe, tromsø. it is also known as noahytten, bunnhuset and laksehytten. holmboehytten 74ø (74°27.1´n 21°41.9´w). hut about 4 km se of giesecke bjerg, southern a.p. olsen land. it was built for salmon fishing in july 1932 for f.a.w. holmboe, tromsø, and is also known as bjørnnesstua and giskehuset. holmenbukta 74ø (74°05.7´n 21°11.5´w). small bay on the se coast of clavering ø, west of eskimovig, and east of eskimohavna. so named by richter (1934) in the archaeological report of the nsiu expeditions, because of the several small islands (= holme) guarding a small harbour. glob (1946) used holmevig for the same feature. holmeslethuset 74ø (74°40.1´n 20°13.9´w). norwegian hunting hut 3 km west of kap schumacher, nw wollaston forland, built in august 1932 by sigurd tolløfsen’s expedition. the name was given for the hunter johannes holmeslet, a member of the expedition. it was also known as kap schumacherhytten. (holmset huset, holme slet.) holmesø 73ø-365 (73°46.3´n 24°50.9´w). small lake in brogetdal, strindberg land, named by hans r. katz during lauge koch’s 1948–49 expeditions for the small island (holm) in the lake. holmevig – see holmenbukta. holmsnes 75ø (75°24.8´n 21°11.3´w). norwegian hunting hut built in august 1932 on the north side of ardencaple fjord for john giæver’s expedition. it had originally been called berglann after an editor of that name in bodø, norway. giæver subsequently renamed it holmsnes after johan holm [b. 1910], a norwegian telegraphist stationed at myggbukta between 1932 and 1936. the hut has also been known as barth-hytta and halsneshytta. (holm nes hytta.) holmneset 72ø (72°43.2´n 21°53.0´w). small peninsula on the east coast of southern geographical society ø. so named on nsiu maps of lacmann (1937), after johan holm (see holmsnes). holstad 71ø (71°45.9´n 22°31.8´w). norwegian hunting hut built in august 1931 for the møre expedition on the west side of wegener halvø. it was named after adolf hoel [1879–1964] (see hoelsbo). the hut has also been known as brown-stua. home forland 73ø-13 (73°50.0´n 20°35.0´w; maps 2, 4). ne part of hold with hope, bounded by tobias dal and tværelv. described by william scoresby jr. in 1822 as a bold and pic tur esque foreland, he named the area home’s foreland after sir everard home [1756–1832], professor of anatomy and surgery at the college of surgeons from 1804 to 1813. scoresby had met home at the house of sir joseph banks in 1815. (home foreland, homes vorland, holmes foreland, home-forland.) homerton 72ø (72°04.3´n 25°10.1´w; map 5). snow dome reaching an altitude of 2360 m west of the head of cavendish gletscher, stauning alper. named by the 1963 cambridge university expedition which climbed to within 30 m of the summit on 20 august 1963. hondal 81ø (81°19.0´n 14°04.1´w). valley in nw kilen, kronprins christian land, perhaps named after the 3-wheeled honda motorcycles used on the 1985 expedition. the name is found on a coloured geological map of kilen printed in 1991 (pedersen 1991). horsedal 71ø-165 (71°50.8´n 23°18.2´w; map 4). valley on the north side of ørsted dal, scoresby land. named by hans stauber during lauge koch’s 1936–38 expeditions following a suggestion by ib poulsen, who travelled this route to antarctic havn with horses (icelandic ponies) in the summer of 1937. horsens fjord 70ø-237 (70°47.8´n 21°45.8´w; map 4). fjord on the east coast of south liverpool land. so named by laurits bruhn during the 1931–34 treårsekspeditionen after the fjord of the same name on the east coast of jylland, denmark. horsnæs fangststation 74ø (74°27.9´n 20°37.9´w). danish hunting station built by nanok in the summer of 1945 on the west side of the river draining into zackenberg bugt, western wollaston for land. this was the proposed original name, as the finances to build the station were offered by the danish newspaper horsens folkeblad. when the funds failed to arrive, the name was changed to zackenberg. horva 73ø (73°02.9´n 23°10.1´w). stream on the north side of geographical society ø, so named on the nsiu (1932a) map. possibly a derivation from the norwegian dialect word for a sea-monster. hospital bugt 70ø (70°29.0´n 21°58.6´w). bay below the hospital at scoresbysund [ittoqqortoormiit], southern liverpool land. the name is used in the ‘den grønlandske havnelods’ (kms 1990). hoved-bræen 70ø (70°18.0´n 29°24.0´w). name occasionally used by carl ryder’s 1891–92 expedition for vestfjord gletscher at the head of vestfjord. hovedet 72ø-203 (72°13.9´n 23°46.4´w; maps 4, 5; see also fig. 198 66). peninsula east of the mouth of noret, north scoresby land. named by prospecting teams associated with lauge koch’s 1948– 49 expedition, for the shape (hoved = head). hovgaard ø 79ø-2, 80ø-2a (79°54.0´n 18°30.0´w; maps 1, 4). island north of nioghalvfjerdsfjorden. named by the 1906–08 danmark-ekspeditionen as hovgaards ø, after andreas peter hovgaard [1853–1910], a danish naval officer and polar explorer. hovgaard took part in the 1881 vega expedition through the ne passage and around asia, and was leader of the 1882–83 dijmphna expedition to the kara sea. hudson land 73ø-25 (73°53.0´n 23°18.0´w; maps 2, 4). land area bounded by loch fyne, moskusoksefjord and promenadedal. the name was adopted by a.g. nathorst in 1899, probably from a british chart. it commemorates henry hudson [d. 1611], who had made an early sighting of east greenland in 1607, a voyage during which he is sometimes said to have discovered jan mayen. the name has been variously applied to larger areas, sometimes including all of the present gauss halvø and hold with hope. hugershoff-fjellet 72ø (72°57.4´n 24°04.7´w). mountain about 1300 m high on western geographical society ø. used only on nsiu maps (lacmann 1937), the name was given for reinhard hugers hoff [b. 1882], a german engineer who made significant contributions to the development of photogrammetric techniques. huggeblokken 76ø-103 (76°46.1´n 18°43.1´w). small skerry in the inner nw part of danmark havn, southern germania land. so named during the 1906–08 danmark-ekspeditionen by charles poulsen. in his published diary (poulsen 1991) he relates that to avoid waking sleeping members of the expedition on their ship the danmark he had rowed out to the skerry to chop up (= hugge) food for the dogs ashore. unfortunately, the dogs saw what he was doing and swam out to the skerry for an early meal (huggeblok = chopping block). hugin 74ø-289 (74°53.8´n 21°27.0´w). mountain 1100 m high on the east side of odin dal, th. thomsen land. the name originated from the wintering party at kulhus during the 1931–34 treårs ekspeditionen. hugin and munin were odin’s two ravens in old nordic mythology, who every morning flew from his shoulder, returning to tell him what was happening in the world. hugin sø 70ø-404 (70°46.1´n 24°05.7´w). small lake in the heden area of sw jameson land. named during the 1967–72 ggu scoresby sund expeditions by svend funder, for its association with the nearby rivers fegin elv and lodin elv, whose names derive from old nordic mythology. see also hugin. hühnerberg-gletscher 74ø (74°28.6´n 19°19.2´w). probably the glacier at the innermost end of gåsedal on the sw side of hühner bjerg, wollaston forland. named by karl koldewey’s 1869–70 expedition, it was briefly studied during the ascent of hühnerbjerg on 11 april 1870. the name is only found on a drawing (verein für die deutsche nordpolarfahrt in bremen 1873–74). see also hühnerbjerg. hühnerbjerg 74ø-5 (74°29.8´n 19°20.5´w). mountain 630 m high in wollaston forland. named by karl koldewey’s 1869–70 expedition as hühnerberg, probably after the ptarmigan (german: hühn = hen). the mountain was climbed by a koldewey party including ralph copeland. the second ascent was made by augustine courtauld on 21 july 1926, who recovered the message left by copeland. andreas vischer noted many ptarmigan during an ascent and remarked on the suitability of the name. it approximately corresponds to william scoresby jr.’s cape beaufoy. (hühner berg.) huitfeldt bjerg 73ø-340 (73°22.2´n 22°14.5´w). mountain in the southern giesecke bjerge. the name was proposed by the place name committee in 1939 to replace suggestions by wolf maync and andreas vischer. it commemorates arild huitfeldt [1546– 1609], a danish historian and nobleman. bessfjellet has also been used. (huitfeldts bjerg.) huledal 71ø-306 (71°33.5´n 24°39.5´w; map 4). valley in karstryggen, west of schuchert flod, at the eastern flank of the stauning alper. named for the large caves (hule = cave) in the limestones of karstryggen by enrico kempter during lauge koch’s 1956–58 expeditions. hulelv [quppaalakajik kangitteq] 70ø-183 (70°33.4´n 22°24.4´w). river in south liverpool land draining west to hurry inlet. so named by laurits bruhn during the 1931–34 treårseks peditionen for the shape of the valley it occupies (hul = hollow, hole). hulesøen 76ø-234 (76°47.2´n 18°44.5´w). lake close to the entrance of gnipahulen, nw of danmark havn, southern ger mania land. so named by the 1906–08 danmark-ekspeditionen because of its proximity to the ice cave gnipahulen (hule = cave). (hule lake.) hullet 71ø-354 (71°13.1´n 27°49.5´w; map 4). ice-dammed lake at the north margin of eielson gletscher, situated in a depression (hullet = the hole). named by the 1963 geodætisk institut expedition. humboldt 73ø (73°06.6´n 23°00.0´w). name commonly used for the norwegian hunting station at kap humboldt, ymer ø, built by arktisk næringsdrift in 1929. see also kjelbotn. (kap humboldt fangststation.) hundedal 72ø-160 (72°11.4´n 22°14.2´w). valley on se traill ø, south of mountnorris fjord, draining into gåsebugt. so named during lauge koch’s 1936–38 expeditions by hans peter schaub (hund = dog). hundehushytten 75ø-96 (75°50.2´n 19°40.2´w). danish hunting hut at the mouth of sønderelv, 13 km north of haystack. built in may 1931 by nanok, and rebuilt in 1932 (hundehus = dog house, dog kennel). it has also been known as terrassehytten. (hunde huset.) hundeklemmen 72ø-176 (72°55.5´n 22°26.9´w). valley on ne geographical society ø. the name was one of a group of names given by the place name committee in 1939, and is said to derive from a danish place name. boykowdalen has also been used. hurry inlet [kangerterajiva] 70ø-148 (70°36.0´n 22°31.0´w; maps 3, 4). fjord between south liverpool land and jameson land. this long fjord was named by william scoresby jr. in 1822 as hurry’s inlet out of respect to mr nicholas hurry, managing-owner of his ship, the baffin. scoresby believed it to be a channel joining up with the present carlsberg fjord and making liverpool land an island (fig. 3). ryder (1895) found that it was a fjord and not a sound. the form hurry fjord often appears on danish maps, al though the original usage hurry inlet is that officially approved. (hurrys inlet, hurry-inlet, hurry bugt, hurry’s einbucht, hurry fjorden.) hurtigrute-tal 73ø (73°45.6´n 23°34.8´w). original name used by heinrich bütler during lauge koch’s 1936–38 expeditions for the present genvejsdalen in moskusokselandet, southern hudson land (e.g. rittmann 1940). the valley was thought to provide a fast (= hurtig) and easy route to the interior of hudson land, but there proved to be a steep ravine in its upper part. husblokken 74ø (74°29.1´n 20°30.9´w). minor locality ne of zackenberg forskningsstation. the name has been used by visiting scientists. husbukta 72ø (72°49.7´n 22°52.5´w). name used for the bay on the south side of geographical society ø where the 1929 nsiu expedition unloaded material for eight hunting huts for arktisk nær ings drift. the name has been used as a reference locality by norwegian and danish botanists. (husbugt.) huselv 73ø-182 (73°30.0´n 21°32.9´w). stream in southern hold with hope flowing into mackenzie bugt. it appears on the nsiu map (nsiu 1932a; fig. 13) as huselva, and was presumably named so because it flows close to myggbukta radio station. huttetu 73ø (73°38.7´n 24°03.9´w). norwegian hunting hut in west ern gauss halvø 5 km north of sydvestpynten, built by john 199 giæver and otto johnsen for arktisk næringsdrift in 1930. ‘huttetu’ is a norwegian expression for unpleasant cold conditions, which prevailed while building the house. it has also been known as sydvestpynten. (huttetuhytten.) hvalpesø 72ø (72°52.6´n 25°06.9´w). lake in the vicinity of ella ø station. the name was used by hammer (1944) for one of the localities where he had collected insects (hvalp = puppy). hvalpynten – see morænepynt. hvalrosbugt [ittoqqortoormiit qinngerajivat] 70ø-302 (70°30.7´n 22°02.1´w). name given during the 1924–25 colonisation expedition for the inner part of rosenvinge bugt, southern liverpool land, because of the numerous walrus which came ashore onto the low gravel beach. in 1924 about 27 were seen on one occasion, and 60 walrus were shot by the greenlanders during the first year of the colony (1925–26). walrus were reported as uncommon here after 1926. an american weather station manned by 20–30 men operated from hvalrosbugt during the war years. (walrus bay.) hvalrosodden 76ø-29 (76°54.6´n 20°06.3´w; map 4). peninsula on the north coast of dove bugt, sw germania land, at the mouth of lakseelven. so named by the 1906–08 danmark-ekspeditionen because they shot 12 walruses here on one of their first boat journeys in august 1906. walruses commonly came ashore to rest here in the early part of the 20th century, but harassed by danish hunters in the 1930s moved away, and currently come ashore at lille snenæs (fig. 47). (hvalrosnæs, hvalros odde, walrus point, hvalros odde, hvalrosnäs, odden, rostungsoddi.) hvalrosodden 76ø-29a (76°55.0´n 20°06.5´w). danish hunting station on the north coast of dove bugt, sw germania land, at the peninsula of the same name. the station was established by østgrønlandske fangstkompagni in 1919, and taken over by nanok in 1929. improvements and a radio station were added in 1932. it was manned in the periods 1919–21, 1932–34, 1938–41 and 1959–60 (p.s. mikkelsen 1994). gunnar andersen died in april 1933 when a party of hunters was trapped by a snowstorm, and was buried beside the station. sirius use and maintain the station, which is regularly visited by parties from danmarkshavn weather station. (hvalrosodden station, odden.) hvalrosskærene 76ø (76°45.5´n 18°47.6´w). small island off wendel pynt, west of danmark havn, southern germania land. according to friis (1909) this was the original name the 1906–08 danmarkekspeditionen had given to the present bådskæret, and arose because alf trolle had shot at a walrus here. hvalrosø 74ø-50 (74°30.8´n 18°45.8´w; maps 2, 4). island south of sabine ø. named by karl koldewey’s 1869–70 expedition as walross insel (fig. 6), because of the numerous walrus seen in the vicinity. one of the østgrønlandske fangstkompagni hunters (lund 1926) noted that the island resembled an enormous walrus lying on the ice, and suggested that this was the reason it received its name. (wallross i., hvalrossön, walrus island.) hvalrosø depotskur 74ø (74°30.4´n 18°46.5´w). small depot hut built on the se side of hvalrosø in the summer of 1921 by øst grønlandske fangstkompagni. it was later used by nanok who moved it to the sw side of the island in 1931. now a ruin. (hvalros ø depotskur.) hvalryggen 77ø-133 (77°06.0´n 23°45.0´w; map 4). hill in north dronning louise land on the south side of britannia sø. so named by the 1952–54 british north greenland expedition because it has a whale-backed crest. hvalsletten 76ø-54 (76°56.5´n 20°06.5´w). extensive plain at the se end of sælsøen, east of hvalrosodden. so named by the 1906–08 danmark-ekspeditionen because the skeleton of a whale was found here several kilometres from the coast, evidently stranded when the sea level was higher than at present. a danish hut built in 1938 a short distance to the north, on the east side of trekroner, has sometimes been known as hvalsletten – see trekronerhytten. hvide ryg 71ø-281 (71°54.3´n 24°10.2´w; map 5). mountain ridge between sirius gletscher and aldebaren gletscher, werner bjerge. it was named by peter bearth and eduard wenk during lauge koch’s 1953–54 expeditions for the light-coloured nepheline syenite rocks. hvide støvhorn 72ø-425 (72°47.9´n 26°53.0´w; map 4). mountain about 2000 m high in gletscherland, on the south side of dickson fjord, named during the 1931–34 treårsekspeditionen by eugène wegmann as white staubhorn. origin uncertain, but see røde støv horn. (weisses staubhorn.) hvidbjerg 72ø-388 (72°01.9´n 23°20.0´w). mountain 974 m high on the east side of majdal, northern scoresby land, partly formed of light-coloured syenite. named by hans kapp during lauge koch’s 1957–58 expeditions. hvidbjørn nunatakker 73ø-592 (73°37.8´n 29°43.3´w; map 4). extensive nunatak group between evers gletscher and hamberg gletscher. this nunatak region was first partly explored by arne høygaard and martin mehren in 1931, and was mapped and named by lauge koch during flights in 1932 on the 1931–34 treårsekspeditionen. it was named after the naval inspection ship hvidbjørn, which had assisted koch’s expedition in 1932. hvidefirn 71ø-284 (71°55.5´n 23°55.8´w). glacier in the southern werner bjerge draining nw along the west flank of mågeborg. named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk. hvidefjeld [apuseeq] 70ø-202 (70°32.0´n 21°44.2´w). ice cap 730 m high in south liverpool land, ne of the town of scoresbysund. the name came into use during the 1924–25 colonisation expedition (e. mikkelsen 1927). french expeditions used the name dôme charcot. (white fjeld.) hvidefjeld 73ø-357 (73°52.9´n 24°43.0´w). mountain over 2000 m fig. 47. walrus (hvalros) off lille snenæs, germania land. photo: jakob lautrup. 200 high in east-central strindberg land, on the north side of rævedal. named during lauge koch’s 1948–49 expeditions by hans r. katz. hvidevæggen 73ø-629 (73°18.7´n 25°38.3´w; map 4; fig. 35). cliff in se andrée land formed by white limestones, named by eugène wegmann during the 1931–34 treårs eks peditionen as white wall. hvidhoved 73ø-679 (73°34.3´n 26°45.9´w). mountain about 2100 m high in western andrée land, north of kalvedal. named by john haller following explorations during lauge koch’s 1949–51 expeditions, for its large, rounded summit ice cap. hynes-hytten – see kap hynæs. hyolithuskløft 73ø-565 (73°31.7´n 24°44.2´w). small ravine in eastern andrée land, nw of kap weber, draining into geolog fjord. named by christian poulsen during lauge koch’s 1929 expedition as hyolithus creek for the finds of numerous fossil hyo li thids. the position of the ravine is incorrectly placed on official place names maps, and published on the geodætisk institut 1:250 000 scale, topographic map. hyttebugt 70ø-245 (70°55.9´n 21°40.4´w). bay on the sw side of kap greg, liverpool land, close to a hunting hut established by scoresbysund municipality. the name was introduced by helge g. backlund in 1935. hyænen 73ø (73°07.8´n 28°32.6´w ). name used in a climbing report by buess (1953) for a summit north of petermann bjerg and sw of kalifbjerg, western frænkel land (hyænen = the hyena). häsi bjerge 72ø-460 (72°14.3´n 27°14.2´w; maps 3, 4). moun tainous region west of violin gletscher. the name was used by eugène wegmann during the 1931–34 treårsekspeditionen, and derives from ‘häsi’, a swiss dialect word for many hares. hætten 76ø-221 (76°58.9´n 20°26.7´w). mountain north of mørkefjord station, between fuglenæbsfjeldet and brystet, daniel bruun land. named by the 1938–39 mørkefjord expedition, presumably for the shape (hætten = the hat, the hood). hödgletscher 74ø-382 (74°22.3´n 21°00.0´w). glacier on north clavering ø draining ne to skilledal. named on the nsiu maps of lacmann (1937) in the form hodbreen, after høder of old nordic mythology, who was lured into killing his twin brother balder. the name was not approved for general usage until 1950. högbom bjerg 73ø-68 (73°36.8´n 22°45.1´w; map 4). mountain 1297 m high on the north side of moskusoksefjord, southern hudson land. named during lauge koch’s 1929–30 expeditions by helge g. backlund, after arvid gustaf högbom [1857–1940], a swedish igneous and metamorphic petrologist, who was professor at the university of uppsala from 1896 to 1922. he was succeeded as professor by backlund. (mt. högbom, høgboms bjerg, högbom berg, högbomfjellet.) högspids 72ø (72°09.0´n 25°16.0´w; map 5). peak about 2100 m high on the south side of vikingbræ, north stauning alper, climbed by hermann huber’s 1968 expedition. (hogspids.) højedal 72ø-433 (72°24.7´n 26°28.4´w; map 4). high hanging valley west of the head of forsblad fjord, sw lyell land. named during the 1931–34 treårsekspeditionen by ove simonsen (høj = high). højkæret 74ø (74°28.1´n 20°38.5´w). vegetated area north of zacken berg basen. the name is used by visiting scientists to zacken berg forskningsstation. højnålen 74ø-73 (74°21.5´n 21°02.4´w). mountain 1512 m high on clavering ø, named by karl koldewey’s 1869–70 expedition as hohe nadel, possibly after a mountain of similar name in austria (hohe nadel = high needle). there is some uncertainty as to the original positions of this mountain and ortlerspids according to seidenfaden (1931), but højnålen is usually placed west of skille gletscher. this was the peak climbed by kaare rodahl in 1939. (mt. hohe nadel, hohe nadeln, hohe nadil, high needle.) højsletten 73ø-333 (73°24.2´n 22°34.6´w). flat-topped mountain on gauss halvø, west of giesecke bjerge. the name was proposed by the place name committee in 1939 to replace suggestions by wolf maync and andreas vischer (højsletten = high plain). hønsetarmen 76ø-96 (76°46.2´n 18°28.1´w). small n–s-trending fjord with marked width variations east of danmark havn, se germania land. so named by the 1906–08 danmark-ekspedi tionen, probably for finds of the plant cerastrium, known as ‘hønsetarm’ in danish (chickweed) because in europe it is eaten by chickens (j. løve, personal communication 2010). høst havn 69ø-29 (69°14.7´n 24°48.0´w). name proposed by ejnar mikkelsen during the 1932 second scoresby sund expedition for a small bay on the north side of barclay bugt. it was given for oluf høst [1884–1966], a prominent danish artist, who had helped to finance the expedition. (høsts havn.) høstakken 71ø-262 (71°58.4´n 24°16.3´w; map 5). mountain about 1100 m high in the werner bjerge, on the summit ridge of malmbjerg. named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk for the shape (høstakken = hay stack). høstakken 70ø (70°28.0´n 26°48.0´w). point on the south side of føhnfjord, about 7 km west of falkepynt. the name is only used in helge vedel’s diaries of carl ryder’s 1891–92 expedition (gulløv 1991; j. løve, personal communication 2010). høygaardbreen 73ø (73°53.5´n 22°36.1´w). glacier in the nørlund alper, north hudson land, corresponding to the present fellen berg gletscher. used only on nsiu maps (lacmann 1937), and named after arne høygaard [b. 1906], a norwegian who made a crossing of the inland ice from west to east with martin mehren in 1931. he wintered in ammassalik in 1936–37. (höygaardbreen.) håbets dal 72ø-379 (72°01´n 23°31´w). valley in north scoresby land, draining south into kolledalen. named by hans kapp during lauge koch’s 1957–58 expeditions (håb = hope). håkampen 73ø (73°26.4´n 22°11.7´w). mountain 1250 m high in the giesecke bjerge, corresponding to the present suhm bjerg. the name was used on the nsiu (1932a) map (håkamp = high knoll). håkonshytta 74ø (74°47.0´n 20°33.2´w). norwegian hunting hut on sw kuhn ø, built in august 1932 for sigurd tolløfsen’s expedition, and named after haakon karlsen, one of the expedition hunters. it is now a ruin. the name is occasionally used as a reference locality in scientific reports (e.g. donovan 1964; koch 1955). (håkons hut, haakonshytta, håkonsstua.) i i.p. jacobsen ø 76ø-86 (76°39.7´n 18°36.0´w). island east of lille koldewey. named by the 1906–08 danmark-ekspeditionen as j.p. jacobsens ö, probably after jacob peter jacobsen [1877–1946]. a physicist and hydrographer, and one of the pioneers of danish hydrographical research, jacobsen had advised alf trolle in obtaining hydrographical instruments for the expedition. the letters ‘i’ and ‘j’ are interchangeable in old danish. (j.p. jacobsens island.) ian’s peak 72ø (72°07.2´n 24°55.0´w; map 5). peak 2607 m high north of majorpasset (col major), northern stauning alper. climbed and named by the 2007 smc east greenland expedition. the name was given for an smc climber, ian angel, who died in 2006. ice col 71ø (71°57.5´n 24°28.4´w; map 5). col on the west side of schuchert gletscher, stauning alper, providing a short cut to a northern branch of storgletscher. it was first traversed and named by the 1961 university of bangor expedition. ida elv 73ø (73°52.4´n 22°01.4´w). river in east hudson land, the present suselv. derived from a girl’s name, it may originate with the work of h.g. backlund during lauge koch’s expeditions in the 1930s. idahøj 75ø (75°10.5´n 19°58.3´w). name used by hans frebold in a report on his work during the 1931–34 treårsekspeditionen, for the present negeren, a mountain 252 m high in south hochstetter 201 forland. girl’s name. idivrodej 69ø (69°53.3´n 22°48.8´w). name recorded by sølberg (1980) for a point on the south side of steward ø, a little west of the eastern cape. it was given for the ruins of 15 inuit houses, a locality where geese breed. tuborg & sandell (1999) use ittikortaajik for the same ruin site. idunbreen 74ø (74°19.2´n 20°50.5´w). glacier on central clavering ø draining nw into skillegletscher. so named on nsiu maps of lacmann (1937) after idun, goddess of youth in old nordic mythology. idwal 73ø (73°41.5´n 25°57.5´w). peak 2102 m high on the north side of grejsdalen, andrée land. climbed by the 2007 army boreal zenith expedition. idwal tooth 73ø (73°41.9´n 25°58.0´w). large rock tower 2162 m high on a ridge on the north side of grejsdalen, andrée land. climbed by the 2007 army boreal zenith expedition. igtâjingmit – see ittaajimmit. igterajik – see itterajik. igterajivit 70ø (70°27.5´n 22°21.0´w). original name for the settlement at kap hope, reported in 1933 by johan petersen, first governor of scoresbysund. it translates as ‘the small houses’. this was also the approved name until 1978, when it was changed to igtâjingmit (now spelt ittaajimmit) to comply with the present usage by the inhabitants. (iderajivit, igterajiut.) igtertivâ – see ittertivaa. igtorqortôrmît, igtorqortôrmît ilivnerat, igtorqortôrmît kímut kangertivat, igtorqortôrmît qíngerajivat – see illoqqortoormiit, ittoq qor toormiit ilinnerat, ittoqqortoormiit kimmut kangertivat, ittoq qor toormiit qinngerajivat. ikaasakajiip nuua 71ø-218 (71°05.0´n 25°42.3´w). prominent cape on north milne land. recorded by the 1955 geodætisk institut name registration, the name translates as ‘the bad fjord’s cape’ and was given for its proximity to ikaasakajik [øfjord]. (ikâsakajîp nûa.) ikaasakajik [øfjord] 70ø-5 71ø-41 (71°00.0´n 26°12.0´w). long fjord or sound between renland and milne land. one of the names recorded by the 1955 geodætisk institut name registration, the name translates as ‘the bad sound’, probably a reference to the frequent strong katabatic winds blowing along the fjord. (ikâsakajik.) ikâsakajik, ikâsakajîp nûa – see ikaasakajik, ikaasakajiip nuua. íkauligssat – see ikkaalissat. ikkaalissat [aamarsuit] 70ø-294 (70°27.7´n 22°14.5´w). aban doned coal mine, a very small coastal excavation in south liverpool land, east of aamarsuit nuaat. the name ikalissat was recorded by alfred rosenkrantz as in use by the greenlanders in 1935, and has also been spelt igaliset or igalisat. íkauligssat, now ikkaalissat, was recorded in 1955 as the form used by the older generation, and was noted also to be that used by the inhabitants at kap hope. ildbjerg 72ø-345 (72°12.5´n 22°36.5´w). mountain 820 m high on se traill ø, so named during lauge koch’s 1956–58 expeditions by h.p. heres. origin uncertain (ild = fire). ilddal 72ø-344 (72°13.4´n 22°35.8´w). valley on se traill ø, draining the north flank of ildbjerg. so named during lauge koch’s 1956–58 expeditions by h.p. heres. île de france (from 2004 qeqertaq prins henrik) 77ø-1 (77°43.0´n 17°45.0´w; maps 1, 2, 4). large island east of jøkulbugten. named by the duke of orléans in july 1905, when the french flag was raised and a cairn built (fig. 9). the cairn message was recovered by eigil knuth in 1988. the name may have been inspired by the name of the french cruise ship île de france that the duke of orléans had met while in svalbard in 1904 (barr 2010). in 2004 the name île de france was officially changed to qeqertaq prins henrik to commemorate the 70th birthday of prince henrik of denmark, french husband of queen margrethe ii of demark. however, the original name is so well established in archaeological and other sciencific publications that it will probably continue to be the preferred usage for many years to come. île lieutier 69ø (c. 69°18´n 25°30´w). island on the northern blosse ville kyst, not identifiable with certainty, but probably a nunatak area north of d’aunay bugt. the name is found on jules blosse ville’s 1833 sketch map (fig. 4), and was given for a french naval officer who was navigator on the la lilloise in 1833 (j. løve, personal communication 2009). ilimanángip nunâ – see ilimananngip nunaa. ilimananngip nunaa [kap leslie] 70ø-48 (70°39.2´n 25°16.4´w). east cape of milne land. one of the names recorded by the 1955 geodætisk institut name registration, it is interpreted as ‘rosen krantz’s land’. alfred rosenkrantz, who worked in the kap leslie region in 1926, was known to his greenlandic assistants as ‘ilimanange’, roughly meaning ‘he one does not expect anything from’. on modern official maps the name is now used as an alternative name for the entire island of milne land. (ilimanángip nunâ.) ilimananngip nunaa 70ø-121 (70°56.0´n 22°33.8´w). valley in east jameson land, carrying the river rødelv. the name was recorded by the 1955 geodætisk institut name registration and, like its equivalent in east milne land, derives from the work of alfred rosenkrantz in the region in 1926–27. (ilimanángip nunâ.) ilinnerajiva 70ø-356 (70°05.6´n 22°08.9´w). sledge route through the valley se of kangikajik [kap brewster]. recorded by the 1955 geodætisk institut name registration, the name means ‘the little crossing place’. (ilivnerajiva). ilinnikajia [roma gletscher] 70ø-341 (70°03.0´n 22°43.3´w). glacier on volquaart boon kyst which is used as a sledge route by hunters travelling from scoresbysund or kap tobin southwards to kap dalton on the northern blosseville kyst. recorded by the 1955 geodætisk institut name registration, the name means roughly ‘the little crossing place’. (ilivnikajia.) ilinnikajiip kiammut kangertiva 69ø-48 (69°59.8´n 22°27.5´w). bay or small fjord sw of kap brewster. the name was recorded by the 1955 geodætisk institut name registration, and translates as ‘the sheltered fjord south of ilinnikajiip kiammut nuaa’. (ilivni kajîp kiámut kangertiva.) ilinnikajiip kiammut nuaa [kap russel] 69ø-1 (69°58.7´n 22°24.6´w). cape on the northern blosseville kyst, sw of kap brewster. the name was recorded by the 1955 geodætisk institut name registration, and roughly translates as the ‘cape to the south of ilinnerajiva’. (ilivnikajîp kiámut nûa.) ilittiartiip nuaa [kap wardlaw] 71ø-15 (71°44.2´n 21°54.1´w). cape in ne canning land. the name was recorded by the 1955 geodætisk institut name registration, and translates as ‘the cape at the little crossing place’. (ilivtiartîp nûa.) ilittiartik 71ø-235 (71°36.3´n 22°25.5´w). low col between nathorst fjord and carlsberg fjord, probably identical with itilleq. recorded by the 1955 geodætisk institut name registration, it means ‘the little crossing place’. (ilivtiartik.) ilivnerajiva – see ilinnerajiva. ilivnikajia – see ilinnikajia. ilivnikajîp kiámut kangertiva, ilivnikajîp kiámut nûa – see ilinni kajiip kiammut kangertiva, ilinnikajiip kiammut nuaa. ilivtiartik – see ilittiartik. ilivtiartîp nûa – see ilittiartiip nuaa. illikasiit / itterajivit – see ittaajimmit [kap hope]. illoqqortoormiut [scoresbysund] 70ø-306 (70°29.1´n 21°57.9´w; maps 3, 4). the town of scoresbysund, south liverpool land, founded in 1924. see also scoresbysund. the greenlandic name for the settlement began as igtorqortôrmît, which translates as ‘those that live at the place with one large house’, and refers to the early days of the colony when the priest and the governor both lived in the only large house. with the revision of spelling, the official spelling became in east greenland dialect ittoqqortoormiit; the 202 settlement newspaper recorded the local spelling in 1984 as iddoqordoormiit. ministry for grønland official documents had begun to use the west greenland dialect form, illoqqortoormiut in the 1970s, and this spelling was imposed on official maps in 1995. east greenlanders continue to use the form ittoqqortoormiit (e.g. arke 2003). the population of the town in 1994 was reported as 484, with an additional 40 in outlying settlements. the 2007 population is recorded as 529. (igtorqortôrmît, igdlorqortôrmiut, igdlor kortôrmiut, ittorqortoorme, ittoqqortoormiit.) ímaqa – see immaqa. imeq [holger danske briller] 71ø-60 (71°25.5´n 25°08.6´w). two large lakes in a valley to the north of sydkap. recorded by the 1955 geodætisk institut name registration, the name means ‘fresh water’. one of the lakes has also been called taseq, which means ‘the lake’. imiilaajiva 70ø-301 (70°30.9´n 22°02.8´w). inner, north side of hvalrosbugt, south liverpool land. recorded by the 1955 geo dætisk institut name registration, the name translates roughly as ‘the cove’. (imîlâjiva.) imiilaajiva 70ø-331 (70°27.4´n 21°49.3´w). inner bay of hartz vig, south liverpool land. recorded by the 1955 geodætisk institut name registration, the name can be translated as ‘the bay with the narrow mouth’, or ‘the channel’. (imîlâjiva.) imikaajik 70ø (70°25.1´n 21°58.6´w). name recorded in 1984 as used by inhabitants at scoresbysund for ravnekløft, which they also call ravneskåret. imîlâjiva – see imiilaajiva. immaqa 74ø-222 (74°01.5´n 21°32.0´w). ravine in nw hold with hope, on the north slope of frebold bjerg, through which river 9 flows. named by eigil nielsen during the 1931–34 treårseks peditionen as immacradal, possibly after the ella ø station boat imara, which was wrecked near store finsch in 1936. this word is commonly used by greenlanders as an answer to a question, and translates as ‘perhaps’. (ímaqa.) immikkeertikajik / immikkeertaajik 70ø (70°49.0´n 22°29.3´w). names used by sandell & sandell (1991) and tuborg & sandell (1999) in their description of iniut ruins on the west side of the largest of the fame øer. the names translate as ‘the little island’. immikkeertaa [depotø] 71ø-51 (71°38.6´n 22°30.0´w). island in the south part of nathorst fjord. one of the names recorded by the 1955 geodætisk institut name registration, it means ‘the island’. (ingmikêrtâ.) immikkeertaa [mågetuen] 71ø-203 (71°32.7´n 26°11.2´w). island in nordvestfjord east of nordbugten. the name was recorded by the 1955 geodætisk institut name registration and means ‘the island’, literally ‘that which sits alone’. (ingmíkêrtâ.) immikkeertaaji 71ø-213 (71°14.8´n 25°14.6´w). two small islands in the centre of the mouth of nordvestfjord. recorded by the 1955 geodætisk institut name registration, the name means ‘the two islands’. (ingmíkêrtâje.) immikkeertaata kangertiva [nordbugten] 71ø-36 (71°35.0´n 26°27.2´w). bay on the north side of nordvestfjord at the mouth of frederiksdal. recorded by the 1955 geodætisk institut name registration, the name means ‘immikkeertaa’s bay’, a reference to the nearby island immikkeertaa [mågetuen]. (ingmíkêrtâta kanger tiva.) immikkeerterajii [menander øer] 72ø-23 (72°20.6´n 24°17.4´w; maps 4, 5). island group on the south side of kong oscar fjord, west of the mouth of skeldal. recorded by the 1955 geodætisk institut name registration, the name means ‘the islands’. (ingmí kêrterajê.) immikkeerterajik 70ø-329 (70°25.6´n 21°51.6´w). small island off the coast of south liverpool land, ne of kap tobin. recorded by the 1955 geodætisk institut name registration, the name means ‘the little island’. (ingmíkêrterajik.) immikkeerterajik 70ø-333 (70°25.9´n 21°47i6´w). island in the mouth of hartz vig, se liverpool land. recorded by the 1955 geodætisk institut name registration, the name means ‘the little island’. in 1984 scoresbysund’s newspaper recorded the local spel ling as immikoordaajik. immikkeerterajik kitterpaaq 70ø-162 (70°47.4´n 22°28.1´w). island south of the fame øer group, the southernmost of the islands at the head of hurry inlet. recorded during the 1955 geodætisk institut name registration, the name translates as the ‘outer island’. (ingmíkêrterajik kíterpâq.) immikkeerterajivit [dunholm] 69ø-24 (69°55.0´n 22°40.0´w). small island ne of steward ø, north blosseville kyst. one of the names recorded by the 1955 geodætisk institut name registration, it translates as ‘the not very large island’. (ingmîkêrterajivit.) immikkeerterajivit iliverta [kap pillans] 69ø-3 (69°56.7´n 22°35.3´w). cape on the northern blosseville kyst, sw of kap brewster. one of the names recorded by the 1955 geodætisk institut name registration, it translates as ‘the cape on the inner side of immikkeerterajivit’. (ingmíkêrterajivit iliverta.) immikkeerterajivit kangittiit 70ø-156 (70°50.0´n 22°30.6´w). northernmost of the islands in the fame øer group. one of the names recorded by the 1955 geodætisk institut name registration, it translates as ‘the inner island’. (ingmíkêrterajivit kangigtît.) immikkeerterajivit qeqqartiit 70ø-157 (70°49.0´n 22°29.2´w). middle island of the fame øer group in hurry inlet. the name was recorded by the 1955 geodætisk institut name registration, and means ‘the middle island’. (ingmíkêrterajivit qeqartît, immikkoortu kajik.) immikkeertikajiip ikaasakajia [turner sund] 69ø-21 (69°45.0´n 23°27.0´w). sound west of immikkeertikajik [turner ø]. the name was recorded by the 1955 geodætisk institut name registration, and translates as ‘immikkeertikajik’s little sound’. (ingmíkêr ti kajîp ikâsakajia.) immikkeertikajiit martik [murray ø, reynolds ø] 71ø-9 (71°32.7´n 21°43.2´w). two islands off the coast of north liverpool land. recorded by the 1955 geodætisk institut name registration, the name translates as ‘the two islands’. (ingmíkêr tikajît martik.) immikkeertikajik [rathbone ø] 70ø-221 (70°40.2´n 21°28.0´w). island off the east coast of south liverpool land. recorded by the 1955 geodætisk institut name registration, the name means ‘the little island’. (ingmíkêrtikajik.) immikkeertikajik [turner ø] 69ø-6 (69°42.0´n 23°24.0´w). island on the north blosseville kyst. one of the names recorded by the 1955 geodætisk institut name registration, it translates as ‘the little island’. (ingmíkêrtikajik.) immikkeertikajik 71ø-223 (71°17.0´n 24°59.4´w). island east of sydkap, at the mouth of nordvestfjord. the name was recorded by the 1955 geodætisk institut name registration, and means ‘the little island’. (ingmíkêrtikajik.) immikkeertikajik kiattikajik [trekanten] 71ø-119 (71°16.6´n 21°42.4´w). island in east liverpool land between campbell sund and tværsund. one of the names recorded by the 1955 geodætisk institut name registration. immikkeertikajik uunertertalik [ janus ø] 70ø-239 (70°52.3´n 21°40.0´w). island off the east coast of liverpool land. recorded by the 1955 geodætisk institut name registration, the name means ‘the little island which has something that burns’, a reference to hot springs on the sw side of the island. (ingmíkêrtikajik ûnartertalik.) immikkeertivaqqat 71ø-222 (71°15.7´n 24°55.8´w). two islands and a skerry east of sydkap. one of the names recorded by the 1955 geodætisk institut name registration, it translates as ‘the small islands’. (ingmíkêrtivarqat.) immikkoortilaq 71ø-199 (71°51.5´n 28°54.7´w; fig. 48). pro mi nent high and narrow peninsula of northern hinks land projecting eastwards into nordvestfjord. it is connected to hinks land by a 203 relatively low neck of land. recorded by the 1955 geodætisk institut name registration, the name can be translated as ‘that which looks like an island’. it has also been known as kap basel (ingmíkôrtilaq). imperial college peak 72ø (72°05.7´n 24°46.9´w). minor peak on the north ridge of merchiston tinde, at the head of bersærkerbræ, stauning alper. the name was used by the 1968 queen mary college expedition during their climb of bersærker tinde via this ridge. a cairn from the 1963 imperial college expedition was found here. indelukket 74ø-302 (74°19.6´n 24°42.7´w; map 4). hidden valley in bartholin land, closed off at both east and west ends by glaciers. the name is said to have originated from the wintering party at eskimonæs during the 1931–34 treårsekspeditionen (indelukket = shut in). inderbredningen 76ø-215 (76°15.0´n 21°37.8´w). broad bay at the front of soranerbreen east of rechnitzer land. named by the 1938–39 mørkefjord expedition, possibly by paul gelting during his journey in april 1939 (inderbredning = inner bay). inderdal 72ø-163 (72°29.2´n 22°18.9´w). valley on east traill ø draining the mols bjerge. named during lauge koch’s 1936–38 expeditions by hans p. schaub for its position within the mountains. it appears as binnental on stauber’s (1938) map. (indlands dal.) inderdalen 73ø-334 (73°21.0´n 22°38.5´w). valley on south gauss halvø, draining east into margrethedal. named by wolf maync and andreas vischer during lauge koch’s 1936–38 expeditions. inderfjord 71ø-115 (71°14.5´n 21°54.5´w). fjord in east liverpool land, sw of kap vidar. so named by laurits bruhn during the 1931–34 treårsekspeditionen because of its situation in the interior of a fjord and island complex. inderhytten 72ø (72°24.4´n 26°02.7´w). norwegian hunting hut built in september 1931 about 10 km from the inner end of fors blad fjord by the møre expedition. it was destroyed by an avalanche in the spring of 1976. it had also been known as bjørktun. inderhytten 76ø (76°35.1´n 18°49.5´w). norwegian hunting hut built in september 1938 by the fransk–norsk polarekspedition on a small peninsula in the ne corner of berg fjord. it is also known as bergfjordhytten. inderhytten 77ø-79 (c. 77°05´n 20°48´w). danish hunting hut on the north coast of inner sælsøen, said to have been built by nanok in 1938. officially known as inderhytten, it has also been known as bundhytten, although in fact the hut was never built (p.s. mik kelsen 1994, 2008). the innermost hut in sælsø was midternæs hytten. ingerborgvatnet 72ø (72°42.6´n 21°54.8´w). lake in extreme se geo graphical society ø. so named on nsiu maps of lacmann (1937), after ingeborg leuch elieson [b. 1884], wife of werner werenskiold. see also werenskioldflya. ingers vig 75ø (75°59.8´n 20°53.0´w). name used for a bay on the north side of bessel fjord by poulsen (1991, p. 191). it may have been named after inger martie thostrup [b. 1884], sister of christian thostrup (j. løve, personal communication 2009). ingmíkêrtâ – see immikkeertaa. ingmíkêrtâje – see immikkeertaaji. ingmíkêrtâta kangertiva – see immikkeertaata kangertiva. ingmíkêrterajê – see immikkeerterajii. ingmíkêrterajik, ingmíkêrterajik kíterpâq – see immikkeerterajik, immikkeerterajik kitterpaaq. ingmíkêrterajivit iliverta, ingmíkêrterajivit kangigtît, ingmíkêrterajivit qeqartît – see immikkeerterajivit iliverta, immikkeerterajivit kang it tiit, immikkeerterajivit qeqqartiit. ingmikêrtikajik kiátikajik, ingmíkêrtikajik ûnartertalik – see immik keerikajik kiattikajik, immikkeertikajik uunartertalik. ingmíkêrtikajîp ikâsakajia – see immikkeertikajiip ikaasakajiip. ingmíkêrtikajît martik – see immikkeertikajiit martik. ingmîkêrterajivit – see immikkeerterajivit. ingmíkêrtivarqat – see immikkeertivaqqat. ingmíkôrtilaq – see immikkoortilaq. ingolf fjord 80ø-9 (80°36.0´n 17°00.0´w; maps 1, 4). fjord between amdrup land and holm land, kronprins christian land. named ingolfs fjord by the 1906–08 danmark-ekspedi tionen after the 544-ton schooner ingolf, which had been used for hydrographic investigations in greenland waters in 1879 and 1895, and which andreas peter hovgaard had used on a voyage to the west indies in 1884–85. (ingolfs fjorden.) ingridbugt 72ø-273 (72°51.2´n 24°53.0´w). minor bay south of lemmingbugt in east ella ø. named by john w. cowie during lauge koch’s 1949–54 expedition after queen ingrid [1910– 2000], wife of kong frederik ix of denmark, and only daughter of gustav vi adolf of sweden. the name was also said to be a tribute to the work of ingrid beck, lauge koch’s long-serving secretary. ingridfjellet 72ø (72°55.3´n 23°39.8´w). mountain 1300 m high on west geographical society ø. used only on nsiu maps (lacmann 1937). girl’s name. fig. 48. distinctive peninsula, immikkoortilaq, projecting eastwards in nordvestfjord. the numerous icebergs have calved from daugaardjensen gletscher just beyond the right side of the photograph. the greenlandic name means ‘that which looks like an island’. it has also been called kap basel. 204 ingridhavn 74ø (74°37.5´n 18°43.9´w). norwegian hunting hut on the south side of hansa bugt in east sabine ø, built by the hird expedition in september 1928. the name appears on an nsiu 1930 list, and seems also to have been used for hansa bugt, or for a very small bay near the hut. the hut has also been known as hansabugthuset. the german meteorological station established nearby by ‘operation holzauge’ in august 1942 was bombed by the us air force in may 1943. (ingrid havn, ingrid-hamn.) ingstadhalvøya 72ø (72°43.8´n 22°04.5´w; fig. 14). peninsula on se geographical society ø. the name is only found on lacmann’s (1937) maps, and was given for helge ingstad [1899–2001], a norwegian lawyer, author and hunter, who was sysselmand (= governor) of eirik raudes land in 1933–34. inkabjerg 73ø-675 (73°38.2'n 26°15.2'w). mountain in central andrée land, on the north side of grejsdal. named by john haller following explorations during lauge koch’s 1949–51 expeditions, and named for the tobacco-brown colour said to be typical of the inca civilisation. it was first climbed by john haller in 1950. inland ice – translation of the danish designation inlandsis for the major ice cap covering central greenland, and the conventional spelling in publications in english (see weidick 1967). indlandsdal 72ø (72°29.6´n 22°19.9´w). valley west of æbeltoft vig that drains south into begtrup vig. the name was used by schaub (1942a, b). indlandsis – the major ice cap covering central greenland, the second largest ice cap in the world. innakajik [kap stewart] 70ø-281 (70°26.6´n 22°38.2´w). cape in se jameson land. one of the names recorded by the 1955 geo dætisk institut name registration, it means ‘the little slope’. (ivnakajik, ideridek.) inner bay 74ø (74°06´n 21°52´w). name occasionally used for the embayment south of clavering ø, the present godthåb golf. inoceramus elv 74ø (74°16.2´n 20°33.4´w). minor river on east clav ering ø, draining north into storstrømmen. the name was used by maync (1949), and was given for finds of fossil inoceramus during lauge koch’s 1936–38 expeditions. ingstadhalvøya 72ø (72°45.0´n 22°15.0´w). se peninsula of geo graphical society ø. so named on nsiu maps of lacmann (1937), after helge ingstad [1899–2001], norwegian lawyer, polar traveller and writer. from 1926 to 1930 he was a hunter in canada, and in 1932–33 led a hunting expedition to east greenland where he was governor (sysselmand) of eirik raudes land. ingstadheimen 71ø (71°37.8´n 22°59.8´w). norwegian hunting hut built in the autumn of 1932 by helge ingstad and normann andersen about 3 km from the inner end of fleming fjord. the name heimen was originally used by ingstad (1937). see also ingstadhalvøya. inner sanctum 71ø (71°58.8´n 25°15.2´w; map 5). inner branch of essemmceebrae, on the south side of sefström gletscher, stauning alper. the name was used by the 1998 scottish mountaineering club expedition. inugsukajik – see inussukajik. inussukajik 70ø-318a (70°27.8´n 21°53.1´w). low hill se of scoresbysund, south liverpool land. the name was recorded by the 1955 geodætisk institut name registration, and means ‘the little cairn’, a reference to a cairn on the summit. (inugsukajik.) inverarnan 72ø-354 (72°01.5´n 25°22.0´w; map 5). mountain with twin summits 2035 m high east of dammen, stauning alper, apparently very close to the mountain metacarpel. the mountain was first climbed by malcolm slesser’s 1958 expedition, and was named after the inverarnan hotel, centre of scottish climbing. invertebrae 72ø (72°09.1´n 25°07.3´w). minor glacier in the stau ning alper south of vikingebræ. ipaqqiarpik 70ø-288 (70°29´n 22°17´w). hillside ne of ittaajim mit [kap hope], sw liverpool land. one of the names recorded by the 1955 geodætisk institut name registration, it translates as ‘where one gathers wall-moss’. (iparqiarpik.) iparqiarpik – see ipaqqiarpik. irene ø 73ø (c. 73°38´n 20°10´w). small island 10 m high and 100 m across off the coast of hold with hope, about 6 km se of knudshoved. the name is used in den grønlandske lods (1968), and is said to have been given by l.m. coulet-svendsen, the first mate on the gustav holm in 1930. girl’s name. irisgletscher 74ø-136 (74°13.8´n 23°23.8´w; map 4). minor southern branch of wordie gletscher, named by lauge koch’s 1929–30 expedition in the form iris glacier, because of the variable colours resembling the iris of the eye. isar-passet 71ø (71°48.1´n 24°57.9´w; map 5). high pass on the sw side of roslin gletscher in the stauning alper, connecting with a branch of mars gletscher. so named by karl herligkoffer’s 1966 expedition after the bavarian river isar. isboksen 73ø (73°17.7´n 24°26.0´w). norwegian hunting hut built in october 1929 by arktisk næringsdrift on the north side of dusén fjord (isboksen = icebox, freezer). it was originally known as devold hytta. isbrosund 70ø-241 (70°52´n 21°45´w). sound between janus ø and the east coast of liverpool land. the name was given by helge g. backlund because the sound was bridged by winter ice during his explorations in 1933 (isbro = ice bridge). isdal 74ø-149 (74°23´n 20°14´w). valley in west wollaston for land, so named during the 1931–34 treårsekspeditionen by hans frebold (isdal = ice valley). isfjeldsund 76ø-285 (76°31.2´n 21°21.0´w; map 4). sound be tween edvard ø and carl heger ø, western dove bugt, where icebergs from bredebræ accumulate. named by the 1938–39 mørkefjord expedition. (isfjældsund.) isfjord 73ø-522 (73°21.7´n 27°00.0´w; maps 3, 4). fjord between andrée land and frænkel land, named by a.g. nathorst’s 1899 expedition as isfjorden because the icebergs encountered were of such colossal dimensions that the ship could not proceed farther than the mouth of the fjord. the head of the fjord was first reached in 1931 by the louise boyd expedition on the veslekari, and later the same summer by the nsiu expedition with the polarbjørn. (ice fjord.) isfjord 77ø (77°48.0´n 20°00.0´w). name used on christian poulsen’s map (1991) for the present orléans sund, in his diary of the 1906–08 danmark-ekspeditionen, presumably because of the presence of ice (j. løve, personal communication 2009). isfjord 70ø (70°28.5´n 28°38.0´w). name used for vestfjord, the fjord extending westwards from rødefjord, in the 1891–92 diaries of helge vedal (gulløv 1991). the name recorded the abundant icebergs calved from vestfjord gletscher. (isfjorden). isfjordhytten – see lille stu. isikajia 71ø-232 (71°08.3´n 22°34.5´w). hill on the floor of the upper part of klitdal, between liverpool land and jameson land. one of the names recorded by the 1955 geodætisk institut name registration, it means ‘the little spire’. isikajia 71ø-239 (71°29.0´n 21°45.5´w). small peninsula in ne liverpool land. one of the names recorded by the 1955 geodætisk institut name registration, it means ‘the little spire’. isjomfruen 70ø-271 (70°03.4´n 23°08.7´w). mountain 1636 m high on volquaart boon kyst. named during the 1931–34 treårs eks peditionen by laurits bruhn for its solitary splendour (= ice maiden). it was first climbed by the 1934 bonzi expedition and given the name punta umberto balestrieri. iskap 80ø-126 (81°07.5´n 12°34.8´w; maps 1, 4). cape on the east coast of kilen, kronprins christian land. the original placement between latitudes 80° and 81°n is due to inaccurate topographic maps. it has also been called iver pynt. islantit [parker øer] 70ø-227 (70°43.4´n 21°29.8´w). small islands east of kap høegh, south liverpool land. recorded by the 1955 geodætisk institut name registration, the islands were so 205 called by the greenlanders because they were situated ‘far off the coast in the direction of iceland’. isle de philippe 77ø (77°43.0´n 17°45.0´w). name occasionally used by the 1906–08 danmark-ekspeditionen for île de france (from 2004 qeqertaq prins henrik), of which the se cape is kap philippe. islington 72ø (72°04.9´n 24°48.3´w). mountain 2400 m high at the head of bersærkerbræ, north stauning alper, the present merchi ston tinde. this name was used by the 1963 imperial college expedition, which made the second ascent, and was given for the london district of islington. the mountain was first climbed by malcolm slesser’s 1958 expedition. ismarken 72ø-432 (72°28.7´n 26°45.6´w; maps 3, 4). ice sheet covering the plateau between wahlenberg gletscher and violin gletscher, western lyell land. named by ove simonsen during the 1931–34 treårsekspeditionen (ismark = ice field). ismågen 69ø-39 (69°03.5´n 29°57.0´w). locality in the nw watkins bjerge, where three ivory gulls were seen, and initially located on maps as the ivory gulls (courtauld 1936). several ivory gull colonies have subsequently been located in this nunatak region. ispassagen 73ø-636 (73°04.2´n 26°33.8´w; map 4). glacier in nw suess land. named during the 1931–34 treårsekspeditionen by ove simonsen. the original description is of a glacier draining to both north and south, which fits nearby borggletscher better than the official location. the two names may have been accidently exchanged. ispynt 70ø-13 (70°26.7´n 28°56.3´w). small peninsula on the north coast of inner vestfjord. named in this form by carl ryder’s 1891–92 expedition, possibly because an ice-filled ravine adjacent to the point was ascended during their sledge journey in may 1892. (is pynt.) issø 70ø-386 (70°07.0´n 28°36.4´w). lake 140 m above sea level at the south side of gåsegletscher, west gåseland. named during lauge koch’s 1958 expedition by eduard wenk. issø [findelen sø] 72ø-288 (72°47.1´n 28°10.0´w; map 4). lake at the north margin of hisinger gletscher at the head of agassiz dal. named during lauge koch’s 1953 expedition by john haller. apart from major features with both greenlandic and danish names, this is one of the few localities in east greenland with two officially approved names. istoppene 73ø-687 (73°31.5´n 26°14.1´w). mountain in the centre of an ice cap in south andrée land. named during lauge koch’s 1949–51 expeditions by john haller. istorvet 70ø-151 (70°55.7´n 22°07.7´w; map 4). large n–s-trending ice cap in central liverpool land. so named by laurits bruhn during the 1931–34 treårsekspeditionen (torv = a square or mar ket place). (istorvet gletscher.) italytinde 72ø (c. 72°10´n 25°10´w). peak 2710 m high in the vikingebræ area of the north stauning alper, climbed and named by g. dionisi’s 1982 expedition. exact location uncertain. itilleq 71ø-89 (71°36.3´n 22°25.5´w). low crossing place in can ning land between nathorst fjord and carlsberg fjord. named during the 1931–34 treårsekspeditionen by arne noe-nygaard as ituidlek, greenlandic for a low area where an umiak (women’s boat) can be carried over land. it is probably identical with ilitti artik. (itivdleq.) itivdleq – see itilleq. ittaajik 70ø (c. 70°28´n 22°23´w). name used by sandell & sandell (1991) for a locality near kap hope, south liverpool land, where inuit ruins were recorded. ittaajimmit [kap hope] 70ø-287 (70°27.5´n 22°20.9´w). settle ment close to kap hope, sw liverpool land, established in 1924 by the colonisation expedition. it was known as igterajivit from 1925 to 1978, when the official name was changed to that used by the inhabitants, igtâjimmit, now ittaajimmit. the name translates as ‘the small houses’. the population in 1970 was a high of 108, reduced to 20 in 2000 and nine in 2005; there were no permanent residents in 2007. the most recent annual statistical reports for greenland use the name itterajivit / illukasiit for the settlement. (ittaaijeme, ittaajimme.) itterajik 70ø-285 (70°29.6´n 22°24.0´w). inuit ruin north of kap hope, on the east side of hurry inlet. recorded by the 1955 geodætisk institut name registration, the name means ‘the little house’. (igterajik.) itterajivit / illukasiit – see ittaajimmit [kap hope]. ittertivaa [kap dalton] 69ø-8 (69°24.7´n 24°04.0´w). cape on the north blosseville kyst. the name was recorded by the 1955 geo dætisk institut name registration. although applied to the cape, the name actually refers to the depot hut, ‘the big house’, built by the 1898–1900 amdrup expedition in the bay on the north side of the cape (see amdrup hytte). (igtertivâ.) ittikajik 71ø (c. 71°14´n 24°36´w). greenlandic name used by tuborg & sandell (1999) for an inuit ruin site near gurreholm on the west coast of jameson land. ittikortaajik 69ø (69°53.3´n 22°50.8´w). peninsula in the se part of steward ø, on the northern blosseville kyst, one of the few locations on the island where it is possible to land, and the site of inuit ruins. the name is used by tuborg & sandell (1999), and means ‘the place with house ruins’. sølberg (1980) used the name idivro dej, that he said the greenlanders called it. this locality is reported to have long been known by hunters from scoresbysund / illoqqortormiut. ittoqqortoormiit – see illoqqortoormiut. ittoqqortoormiit ilinnerat 70ø-201 (70°32.0´n 21°51.2´w). sledge route, or crossing place, between illoqqortoormiut and lillefjord. recorded by the 1955 geodætisk institut name registration, the name translates as ‘the crossing place from illoq qortoor miut’. the local newspaper recorded the spelling iddo qordoormiit ilinnerat in 1984. (igtorqortôrmît ilivnerat.) ittoqqortoormiit kimmut kangertivat [amdrup havn] 70ø-312 (70°28.4´n 21°54.6´w). fjord or harbour east of illoqqortoormiut [scoresbysund]. the name was recorded by the 1955 geodætisk institut name registration, and describes its location, ‘illoqqor toormiut’s eastern fjord’. the local scoresbysund newspaper recorded in 1984 the name endalip kangersiva for this feature. (igtor qor tôrmît kímut kangertivat.) ittoqqortoormiit qinngerajivat [hvalrosbugt] 70ø-302 (70°30.1´n 22°02.1´w). inner part of rosenvinge bugt, south liverpool land, west of the town illoqqortoormiut [scores bysund]. recorded by the 1955 geodætisk institut name registration, the name translates roughly as ‘the bottom of illoqqor toor miut’. (igtorqortôrmît qíngerajivat.) ittorisseq 70ø-284 (70°27.3´n 22°37.0´w). former settlement north of kap stewart. this was one of the original sites chosen by the founders of the scoresbysund colony for hunters settlements. three houses were built in 1924, and ryder’s depot house built here in 1892 was repaired. the site fell into disuse about 1930 due to frequent heavy snow, and subsequently has mainly been visited by hunters from kap hope / ittaajimmiit (sandell & sandell 1991). the name was recorded as ivtorigseq by ejnar mikkelsen in 1925 and johan petersen in 1933. it translates roughly as ‘here there is good turf ’. turf was used to build the traditional greenlandic winter houses. (itoissoq, ivtssorigsek, ittoritteq, ivssorigsed, ivssorig sek.) iuel-brockdorff bjerg 77ø-49 (77°11.3´n 24°50.5´w; fig. 21). nunatak in nw dronning louise land, named during the 1909– 12 alabama expedition as juel-brockdorff ’s nunatak, probably by vilhelm laub. this nunatak region was explored by laub, who had sailed with juel-brockdorff aboard the islands falk to iceland in 1909. niels juel-brockdorff [1878–1964] was a naval officer, from 1904 a first lieutenant and from 1915 a captain (j. løve, person206 al communication 2009). the letters ‘i’ and ‘j’ are interchangeable in old danish. ivar baardsøn gletscher 71ø (71°48.0´n 24°48.2´w; map 5). large glacier in the stauning alper draining se into schuchert dal, the present roslin gletscher. this name was one of a group of names for glaciers given by the place name committee in 1939. it was also the officially approved name from 1939 to 1971, although had only rarely been used on maps (e.g. kempter 1961; cruikshank & col houn 1965). due to some confusion, and the lack of accurate topographic maps, the name roslin gletscher was approved for the same glacier in 1959. roslin gletscher became widely used, and the use of ivar baardsøn gletscher was abandoned in 1971. the original name had commemorated ivar baardsøn, a priest from the bergen region of norway who was sent as bishop to the norse settlements of greenland at the end of the 14th century. he is noted for his description of greenland. (ivar baardsöns gletscher.) iver pynt 81ø (81°07.5´n 12°34.8´w). peninsula on the coast of eastern kilen, kronprins christian land, identical with the approved name iskap. the name is found on a coloured geological map of kilen printed in 1991 (pedersen 1991), and was given for iver p. iversen, who accompanied ejnar mikkelsen during the 1910 search by the alabama expedition for the lost members of the 1906–08 danmark-ekspeditionen. ivingdalen 74ø (74°21.1´n 20°29.1´w). valley on ne clavering ø. used only on nsiu maps (lacmann 1937), the name is derived from old nordic mythology. ivnakajik – see innakajik. ivtoriseq – see ittoriseq. j j.f.b. mountain 72ø (72°53.7´n 27°49.0´w). mountain on the west side of bocksrietdalen, the present hagar bjerg. named by louise boyd during her 1931 expedition in memory of her father john f. boyd, a pioneer in the usa mining industry, whose financial suc cess made possible her series of arctic voyages. j.h.l. vogt’s fjeld – see vogt bjerg. j.l. mowinckel land 73ø-577 (73°51.0´n 28°27.9´w; maps 2–4). mountainous region south of adolf hoel gletscher and west of andrée land. named by arne høygaard and martin mehren during their 1931 expedition as j.l. mowinckels fjell for the nor wegian prime minister [johan ludwig mowinckel, 1870–1943], who had shown interest in their expedition. state contributions covered half the expedition expenses. (i.l. mowinckel land.) j.p. koch fjeld 70ø-106 (70°40.5´n 22°55.6´w; map 3, 4). hill 909 m high, the highest point in southern jameson land. the name was used in the form j.p. koch mountain by alfred rosenkrantz in l. koch (1929a), and was given for johan peter koch [1870–1928]. a danish army officer and explorer, koch took part in g.c. amdrup’s 1898–1900 expedition, the 1906–08 danmark-ekspe ditionen as leader of the cartographic work, and in 1912–13 accompanied by alfred wegener led an expedition across the inland ice. (mt. i.p. koch fjæld, j.p. kochs fjeld.) j.p. jacobsen ø – see i.p. jacobsen ø. jaalspids 72ø (72°07.4´n 24°58.3´w). peak in the stauning alper on the sw ridge of dansketinden. it was climbed by the 1996 scottish mountaineering club expedition. jacks hytte 74ø (74°36.4´n 19°40.7´w). danish hunting hut built by nanok in august 1950 on the west side of brorson halvø, northern wollaston forland. the name commemorates an incident in the winter of 1949–50, when the hunter jack christensen lost two of his toes to frostbite. jackson ø 73ø-14 (73°54.9´n 20°07.6´w; map 2, 4). island ne of hold with hope. named by william scoresby jr. in 1822 as jackson island after thomas jackson of whitby, who had married scores by’s third and youngest sister arabella in 1812. he was also a cousin of scoresby’s. jackson’s son, robert edmund scoresby-jackson, wrote a biography of william scoresby. (jacksons ö, jack son insel, jakson insel, jacksonöya, jacksonøya.) jacksonstua 73ø (73°54.3´n 20°09.6´w). norwegian hunting station on sw jackson ø, built by the hird expedition in 1928. the station was manned only in the periods 1928–29 and 1933–34, and was subsequently occasionally used by hunters (p.s. mikkelsen 2008). it was maintained by sirius, until accidently burnt down in 1981. (jackson hytten, jacksonhytta.) jacksontoppen 73ø (73°55.7´n 20°07.1´w). highest point of jackson ø, 422 m in altitude. the name appears on the nsiu (1932a) map. jacob’s house 69ø (69°54.6´n 22°56.2´w). name used by tuborg & sandell (1999) for one of four hunters houses in a bay on the nw side of steward ø, northern blosseville kyst. the houses are used by hunters from scoresbysund, who regularly overwinter here, and the first house was built in 1971 by jakob sanimuinaq. jägmästeren ø – note that ä is treated as æ in danish, thus jäg mästeren ø is listed after jægersund on page 210. jakhellnsundet 72ø (74°45.0´n 23°00.7´w). narrow sound between kista ø and traill ø in vega sund, corresponding to the present snævringen. the name is used on the nsiu maps of lacmann (1937), and was given for anton jakhelln [1904–1990]. a norwe gian meteorologist and oceanographer, he took part in nsiu expeditions in 1931 and 1932, and was in the antarctic in 1933–34. jakob dal 73ø-343 (73°28.9´n 22°04.1´w). valley in the giesecke bjerge, draining eastwards. it was named by wolf maync and andreas vischer during lauge koch’s 1936–38 expeditions after jakob sanimuinak of scoresbysund, their greenlandic assistant and sledge-driver in 1937 and 1938. (jakobsdal.) jakob kjøde bjerg 74ø-177 (74°08.2´n 26°32.4´w; map 4). large nunatak 1850 m high on the north side of adolf hoel gletscher. named by arne høygaard and martin mehren during their 1931 expedition as jakob kjødes fjell, for one of norway’s largest ship owners [jakob kjøde 1880–1946]. (jacob kjødes fjell, jakob kjødes bjærg.) jakob severin bjerg 71ø-196 (71°13.0´n 23°31.3´w). mountain in central jameson land, south of olympen. the name was one of a group of names given by the place name committee in 1939 to replace proposals by hans stauber. it was given for jakob severin [1691–1753], a danish businessman and nobleman who acquired the trading rights in greenland in 1734, and founded the colonies of christianshåb, frederikshåb and jakobshavn. jakobsbo 72ø (72°02.3´n 24°03.7´w). two huts, built in 1960 by nordisk mineselskab in the interior of deltadal, were known by this name, which commemorates ‘gamle jakob’, a carpenter from minebyen, the lead mine near mestersvig airfield. they were built in connection with bulldozer transport between mestersvig and malmbjerg. (jacobsbo.) jakobsendalen 72ø (72°59.6´n 23°22.0´w). valley on north geo graphical society ø. used only on nsiu maps (lacmann 1937), it was named after anton karl hagbart jakobsen [1874–1983], a nor wegian bank director and politician, who was also a shipowner. jaksla 72ø (72°58.1´n 24°500.1´w). mountain 1330 m high on west geographical society ø. used on the nsiu maps of lacmann (1937), and apparently named for its tooth-like shape. jamesondal 71ø (71°39.0´n 22°45.0´w). name occasionally used for the valley on wegener halvø containing jameson elv. jameson elv 71ø-85 (71°39.0´n 22°45.0´w). river on wegener halvø draining ne into nathorst fjord. named by arne noenygaard during the 1931–34 treårsekspeditionen as jameson river, because of its proximity to jameson land. jameson land 70ø-91 71ø-122 (71°00.0´n 23°15.0´w; maps 3, 4). extensive land area bounded by hall bredning, scoresby sund, hurry inlet and carlsberg fjord, with its northern boundary fixed in 1966 following major paar dal, coloradodal, olympen and 207 passagen at about latitude 71°35'n. named by william scoresby jr. in 1822 as jameson’s land (fig. 3) in token of friendship to robert jameson [1774–1854], professor of natural history at edinburgh from 1804. he became scoresby’s friend and mentor, and introduced him to edinburgh society. jameson contributed the appendix on rock specimens to scoresby’s (1823) narrative. janus ø [immikkeertikajik uunartertalik] 70ø-239 (70°52.3´n 21°40.0´w; map 4). island off the east coast of southern liverpool land, so named during the 1931–34 treårsekspeditionen by lau rits bruhn after janus sørensen who had visited scoresbysund in 1927–28 to build a radio station, and the seismic station of which he was leader. he prepared a map of south liverpool land on the basis of his sledge journeys. japetus bjerg 72ø-136 (72°13.1´n 22°42.7´w; map 4). mountain on south traill ø, nw of drømmebugt. the name came into use during lauge koch’s geological expeditions in the 1930s, and is attributed to helge backlund. the name may have been given for japetus steenstrup, see steenstrup bjerg. jarner plateau 77ø-75 (77°36.1´n 19°24.8´w; map 4). plateau in se stormlandet, north of skærfjorden. so named by david malmquist during the 1931–34 treårsekspeditionen after hakon høeg jarner [1882–1964], the geologist of the 1906–08 danmark-ekspedi tionen. see also kap jarner. jarners hytte 76ø (76°28.5´n 21°41.2´w). name occasionally used for bræfjordhytten, southern lindhard ø, north of the mouth of bræfjord. a danish hunting hut, it was built by nanok in may 1934. now a ruin. see also kap jarner. jarners kulmine 75ø-54 (75°11.5´n 19°59.8´w). coastal coal outcrops in sw hochstetter forland, originally found by julius payer during karl koldewey’s 1869–70 expedition. the locality was relocated by h.h. jarner in may 1908 during the 1906–08 danmark-ekspeditionen, and found again in april 1927 by lauge koch. the name appears to have originated from j.g. jennov and richard bøgvad who visited the locality in 1930, and was perhaps first used in the geological report of frebold (1932). lauge koch laid claim to the ‘mine’ in 1931 on behalf of the danish state, when the godthaab took on board 36 tons of coal here. kulfjeldet has also been used. see also kap jarner. the name jarners kulmine has occasionally been used for the danish station kulhus that was erected at this location. (jarner’s kohlengrube.) jassdal 72ø (72°23.6´n 24°51.5´w). name used in a climbing report by braun (1953) for the present skipperdal in the north stauning alper. braun accompanied erdhart fränkl on his geological explorations during lauge koch’s 1950–51 expedition. see also jasspas. jasspas 72ø-303 (72°02.0´n 23°53.9´w; map 5). col or pass on the south side of aggersborg, south of mesters vig, named by peter bearth and eduard wenk during lauge koch’s 1953–54 expeditions. origin uncertain, but possibly named after jass, a swiss cardgame played with a deck of 36 cards. jeannet bjerg 72ø-111 (72°40.6´n 25°03.7´w; map 4). ice-capped mountain 1800 m high in east lyell land, named during the 1931–34 treårsekspeditionen by eugène wegmann in the form mt. jeannet, for a swiss professor. it was climbed by wegmann’s geological party in july 1933. (jeannets bjerg, mont jeannet.) jelsdal 72ø-114 (72°32.7´n 25°34.2´w; map 4). valley in lyell land draining east into polhem dal. the name was an adaptation by the place name committee of a suggestion by eugène wegmann in 1935 (jezlerdal), rejected because it was thought to be a family name. jelstrupfjellet 72ø (72°57.9´n 23°44.9´w). mountain on west geo graphical society ø. so named on nsiu maps of lacmann (1937) for hans severin jelstrup [1893–1964], a norwegian astronomer who took part in the 1931 and 1932 nsiu expeditions to east greenland, and also expeditions to svalbard. jennovs næse 76ø (76°23.8´n 20°48.6´w). name reported by the 1952–54 british north greenland expedition as in regular use by the personnel at danmarkshavn weather station for the prominent mountain sylen, near ålborghus. it commemorates johannes ger hardt jennov [1886–1980], founder of nanok østgrønlandske fangstkompagni (commonly known as nanok), of which he was director from 1929 to 1976. the name jennovs næse has been used by nanok hunters for a number of other mountains in east green land. jennovshåb 74ø (74°47.7´n 19°50.5´w). danish hunting hut built for nanok in september 1930 about 8 km south of kap maurer, kuhn ø. it was more commonly known as kap maurer hytten. see also jennovs næse. jens munk plateau 71ø-175 (71°28.1´n 23°29.5´w; map 4). plateau in northern jameson land, north of olympen. the name was one of a group of names given by the place name committee in 1939 to replace proposals by hans stauber. it commemorates the admiral jens munk, who was sent out by christian iv of denmark and norway in 1619 to find the nw passage. jensenhytta 73ø (c. 73°42´n 23°48´w). norwegian hunting hut east of kap kolthoff on the south side of moskusoksefjord, built in 1930 by arktisk næringsdrift. it was swept away by an avalanche in 1954. perhaps identical with johnsen-hytta, of which jensenhytta may be a variant. jenssonhøgda 72ø (72°54.6´n 22°15.2´w). range of low hills on east geographical society ø. so named on nsiu maps of lacmann (1937) for the norwegian journalist gunleik jensson [b. 1891], who accompanied the 1929 nsiu expedition to greenland. jernhatten 71ø-251 (71°57.8´n 23°52.1´w; map 5). mountain in the eastern werner bjerge, south of antarctic pas. the name was given by the place name committee in the 1950s (jern = iron). it was climbed by peter bearth in 1953. jernhatten 74ø-284 (74°07.4´n 21°00.0´w). mountain on se clavering ø. the name originated from the wintering party at eskimonæs during the 1931–34 treårsekspeditionen, and records the rusty red weathering colour of the summit rocks. jernvæggen 76ø-224 (76°57.3´n 21°07.4´w). mountain in daniel bruun land, on the north side of inner mørkefjord. named by the 1938–39 mørkefjord expedition, presumably for its appearance (jernvæggen = the iron wall). jezlerdal – see jelsdal. jo-neset 72ø (72°59.0´n 24°33.4´w). name used in an nsiu report (1932c) for a cape on the south side of the mouth of sofia sund, north of svedenborg bjerg, where a hunting hut (jopladsen) was built in september 1930. the cape has also been known as kapp 7. juni. joakimpasset 72ø (72°54.0´n 25°05.8´w). pass on central geograph ical society ø, corresponding to part of tværdal. used only on nsiu maps (lacmann 1937), and named after joakim devold [b. 1908], a norwegian artist who took part in expeditions to east and se greenland in 1931 and 1932. jobjerg 71ø (71°59.0´n 24°55.3´w; map 5). summit 2330 m high on the west side of upper storgletscher, central stauning alper. climbed and named after a living person by the 2007 smc east greenland expedition. joffert-gletscher 72ø (72°01.0´n 24°09.1´w; map 5). northern of three small glaciers between vestre gletscher and mellem glet scher in the northern werner bjerge. the name was used by styger (1951) in a report on a climbing excursion during lauge koch’s 1950 expedition, and was named after japheth, one of the sons of noah. see also ham-gletscher and sem-gletscher. joh. g. guildal ø 76ø-92 (76°42.0´n 18°29.8´w). small island east of kap bismarck. named by the 1906–08 danmark-ekspeditionen as guildals ö, after johan peter samuel goldschmidt guildal [1855–1920], a danish businessman who was on the board of directors of many important danish companies. he made a contribution to the expedition finances. joh. h. andresenfjellet 74ø (74°15.0´n 21°51.0´w). mountain ridge 208 on sw clavering ø, equivalent to the present hallebjergene. used on the nsiu maps of lacmann (1937), and named after johan henrik andresen [b. 1888], norwegian businessman and owner of j.l. tiedemanns tobaksfabrik (an oslo-based norwegian tobacco company). his financial support made possible the aerial photography undertaken on the 1932 nsiu expedition. johan davidsen dal 73ø-93 (73°55.8´n 23°58.6´w). valley in west hud son land draining sw from krumme langsø to waltershausen gletscher. named during the 1931–34 treårseks peditionen by th. johansen after his greenlandic assistant (johan davidsen). moskus dalen has also been used. a large ice-dammed lake periodically forms at the margin of waltershausen gletscher, and when empty the fine-grained silt on the lake-bottom may be lifted by katabatic winds to form large clouds that have been mistaken for volcanic eruptions. norwegian newspapers carried reports of ‘volcanic eruptions’ seen by john giæver and charles swithinbank in august 1952. similar reports in 1931 led directly to the 1932 expedition by sigurd skaun and harald welde. (johan davidsental.) johan ligners älv 74ø (74°28.1´n 20°35.7´w). river flowing into young sund near zackenberg, where johan ligner, a swedish doc tor from örebro, fished for salmon (arctic char) in 1937. the name is only used in munsterhjelm (1937). johan olsen-högda 73ø (73°31.9´n 21°29.0´w). hill ne of mygg bukta station in southern hold with hope. the name occurs on the nsiu map (1932a; fig. 13), and was probably given for johan a. olsen whose 1922 expedition built the first myggbukta radio sta tion. the entire expedition was lost when the anni 1 was crushed in the ice on the way home in 1923. johannes knudsens topper 73ø (73°58.6´n 24°20.7´w). name used by sigurd skaun and harald welde in 1932 for a mountain with two characteristic tops in southern ole rømer land, north of posten. johannesendalen 73ø (72°59.3´n 23°39.8´w). valley on west geogra phical society ø draining north into sofia sund. the name is used only on nsiu maps (lacmann 1937), and was given for sigurd hal v orsen johannesen [1881–1964], a norwegian businessman who was a member of several norwegian ministries connected with whaling and fishing. john phillips dal 72ø-258 (72°59.4´n 22°26.7´w). valley on ne geographical society ø. named during lauge koch’s 1949–50 expedition by desmond t. donovan for john phillips [1800–74], a geologist who became professor of geology at kings college lon don, and the universities of dublin and oxford. he was a nephew of william smith (see william smith dal). johns hytta 76ø (75°58.9´n 21°22.0´w). norwegian hunting hut built in september 1932 by john giæver’s expedition on the north side of bessel fjord, ad. s. jensen land. it was named after john johnsen who helped to build it. now a ruin (1988). (john johnsens hytte, johnshytten.) johnsen-hytta 73ø (c. 73°42´n 23°48´w). norwegian hunting hut east of kap kolthoff at the mouth of moskusoksefjord. erected in november 1930 by arktisk næringsdrift, it was named after the norwegian hunter otto johnsen who helped build it. it has now disappeared. see also otto johnsenvika. the hut has also been known as jensenhytta and kolthoffhytten. (johnsenhytten). johnstrup bjerg 72ø-45 (73°00.9´n 25°21.3´w; map 4). ice-capped mountain about 1860 m high in eastern suess land. named by a.g. nathorst’s 1899 expedition as johnstrups berg after johannes frederik johnstrup [1818–1894], a noted danish geologist who was professor of geology and mineralogy at the mineralogisk muse um, copenhagen, from 1866–1895. he was the first chairman of com missionen for ledelsen af de geologiske og geografiske undersøgelser i grøn land that later became kommissionen for videnskabelig under søgelser i grønland (the commission for scientific research in greenland), and he was also an editor of the journal meddelelser om grønland. (johnstrup mountain.) joinville ø 77ø-10 (77°29´n 19°50´w; map 4). island in the inner part of skærfjorden. named by the duke of orléans in 1905 as i. joinville, possibly for his grandfather’s brother, françois ferdinand philippe d’orléans, prince de joinville [1818–1900]. jomfru gletscher 72ø-521 (72°08´n 27°43´w; map 4). glacier in western nathorst land draining into jomfrudal. named by geoffrey halliday following botanical work during the 1961 leicester university and 1971 northern universities expeditions. (jomfrubræ.) jomfru tidsfordriv fjord 79ø-42 (79°13.0´n 19°42.0´w; map 4). narrow n–s-trending fjord in eastern lambert land. this was one of five names given by the place name committee after dogs used on the 1906–08 danmark-ekspeditionen. ‘jomfru tidsfordriv’ was a lady dog of good repute who knew how to keep the gentleman dogs at a distance. the dog was named after a noted copen hagen character, juliane maria hansen, the daughter of a priest. when jilted by a lieutenant, she took to wandering the streets of copenhagen in a green skirt and large boots, and gave sweets to the children who called after her. jomfru-hytta 72ø (72°43.8´n 22°37.3´w). norwegian hunting hut built in august 1929 by arktisk næringsdrift on se geographical society ø, at the mouth of malia havn. (jomfruen). jomfrubjerg 72ø-472 (72°07.3´n 27°01.2´w). mountain 2210 m high at the confluence of herthadal and jomfrudal, nathorst land. named during lauge koch’s 1954–55 expeditions by hans zweifel. jomfrudal 72ø-437 (72°04.8´n 27°03.2´w; map 4). narrow valley west of violingletscher, named during the 1931–34 treårseks peditionen by ove simonsen because its hidden position meant it was virgin territory (= jomfru). jomfruelv 72ø (72°04´n 27°09´w). name used by geoffrey halliday during the 1961 university of leicester expedition after the river in jomfrudal, west nathorst land. jomfruen 70ø-421 (70°36.8´n 29°27.3´w; map 4). nunatak 1770 m high in the upper part of rolige bræ, north of paul stern land. named by laurent jemelin during the 1967–72 ggu scoresby sund expeditions for its isolation and appearance. jomfrupollen 72ø (72°41.7´n 22°37.9´w; fig. 14). small, nearly en closed bay on the south side of geographical society ø, corresponding to malia havn. the name was used by norwegian hunters as early as 1929, and occurs also on nsiu maps (lacmann 1937); it can be translated as ‘virgin bay’. jomsborg 73ø-660 (73°21.4´n 26°38.3´w; map 4; fig. 49). moun tain 1900 m high in sw andrée land west of renbugten, with a conspicuous, near-vertical, se cliff face rising more than 1300 m from the fjord. the name originated from the 1931–34 treårs ekspeditionen, and was approved at the suggestion of r. spärck. it commemorates the fortress of jomsborg, founded by palnatoke, the hero of the joms vikings saga. jomsborg dal 73ø-692 (73°23.8´n 26°27.3´w). valley in sw andrée land, east of jomsborg on the opposite side of rendalen. named by john haller following explorations during lauge koch’s 1949–51 expeditions. jomsdal was used by john haller, who also used joms gletscher for the glacier draining westwards into the valley. jónsbú 75ø-61 (75°19.2´n 20°23.3´w). norwegian hunting station on the west side of peters bugt, ne of the mouth of ardencaple fjord. it was erected by john giæver’s expedition in 1932, and named jonsbu for john schjelderup giæver [1901–1970], a journalist from 1923–1929, a hunter in east greenland from 1929–1934, and from 1935 secretary of nsiu. he was one of the best known of norwegian hunters, and noted for his many books on hunting and the arctic (e.g. giæver 1930, 1931, 1937, 1939, 1958). the original station, to the ruin of which the name is still officially applied, was burnt down in august 1943 by a patrol from the us ship northland to prevent it from being used by german forces. in 1948 a new jonsbu hunting station was built on the south side of 209 ardencaple fjord se of the mouth of kildedal (75°14.8´n 20°52.6´w). the two huts have sometimes been distinguished as gamle jonsbu and ny jonsbu. the accents on the approved version of the name were added by the place name committee as an aid to correct pronunciation. norsk petersbugt station has also been used. (johnsbu.) jones-fairey spur 71ø (71°56.5´n 25°03.8´w). this is described as the sw spur of a western outlier of sefströmsgipfel that was climbed by the 2001 smc east greenland expedition to reach point jonesfairy (2570 m). it is located in the upper reaches of sefström glet scher. joplassen 72ø (72°59.0´n 24°33.4´w). norwegian hunting hut built in 1929 at kapp 7. juni, the nw point of geographical society ø, which has also been known by norwegians as jo-neset. the hut has also been known as svedenborg and valborghytten, and in recent years has also been called røvballehytten. jordanbukta 74ø (74°09.9´n 22°18.5´w). small bay between kap adam and kap eva, north of jordan hill, equivalent to the present hansen havn. used by norwegian hunters, the name appears on the nsiu (1932a) map. jordanhill 74ø-21 (74°07.6´n 22°19.9´w; maps 2, 4; fig. 15). pro minent landmass 1410 m high at the front of wordie gletscher. so named by douglas clavering in 1823, who climbed to within 200 m of the top, after the residence of his friend james smith. see also kap james. jordanhill is near glasgow, scotland. (jordanhill insel, jordan hill.) jordanhill glacier 74ø (74°15.0´n 23°05.0´w). name used by j.m. wordie in 1926 for the large glacier west of jordanhill now known as wordie gletscher. jordanhill hytta 74ø (74°06.7´n 27°10.9´w). norwegian hunting hut on the east coast of jordanhill, built by arktisk næringsdrift in august 1953 as a replacement for jordanstranda. it is now a ruin. jordanstranda 74ø (74°06.7´n 22°10.9´w). norwegian hunting hut on the east coast of jordan hill, built by the foldvik expedition in september 1927. it was replaced by a new hut known as jordanhill hytta in 1953. (jordan-stranda, jordan.) jordflommen 74ø-311 (74°05.8´n 21°15.4´w). solifluction flow on the east side of østhavn, east of eskimonæs station, south clav ering ø. the name originated from the wintering party at eski monæs during the 1931–34 treårsekspeditionen. detailed un published maps (1:10 000) show two such flows to which the name could be applied, on either side of østelv. jordly 73ø (73°45.5´n 20°59.6´w). danish hunting hut in central tobias dalen, hold with hope, built by nanok in the spring of 1945 (jord = earth). it has also been known as vulkanhytten. jostein 72ø (72°07.5´n 23°28.6´w). hunting hut 15 km nw of kap syenit, east of the mouth of mesters vig. it was built in 1930 by the møre expedition, and named after jostein, the youngest son of odd åmbakk, one of the hunters. it has also been called segldalen, bjørnebu and pictetbjerghytten. (josteinshytte.) jostgletscher 74ø (74°16.2´n 21°12.6´w). name used by mittelholzer (1941) for the present snemarken, central clavering ø, in his report on work during lauge koch’s 1938–39 expeditions. (josts gletscher.) josvakajiip kaporniagaqarpia 70ø (70°21.1´n 28°08.0´w). name sometimes used for the hut at the mouth of hjørnedal, where fønfjord and rødefjord meet. jotunheim 75ø-78 (75°14´n 22°38´w; map 4). ice plateau in western c.h. ostenfeld land. the name originated from the wintering party at kulhus during the 1931–34 treårsekspeditionen, and was given for the region of the same name in south norway. jubilee peak 71ø (71°18.9´n 21°54.3´w). mountain 1048 m high west of stensund, liverpool land, climbed by four members of the 1977 joint services expedition. the name, given to commemorate the 25th jubilee year of queen elizabeth ii’s accession, was reported in several british newspapers. juel-brockdorff ’s nunatak – see iuel-brockdorff bjerg. julekagen 72ø-177 (72°53.5´n 23°04.6´w). mountain range be tween græsdalen and lysedal, geographical society ø. the name was derived from a suggestion by hans stauber during lauge koch’s 1936–38 expeditions, who had compared the mountain to a giant christmas cake (= julekagen). juliasbjerge 71ø (71°59.6´n 24°55.3´w; map 5). peak 2058 m high on the west side of upper storgletscher, central stauning alper. climbed and named after a living person by the 2007 smc east greenland expedition. jûlut dal [ juluut dal] 73ø-638 (73°05.0´n 24°28.8´w; map 4). valley on south ymer ø, draining west to karl jakobsen bugt. named by ove simonsen during the 1931–34 treårsekspeditionen for jørgen petersen, known as jûlut, a greenlander who acted as assistant and dog-driver throughout the expedition and hunted in this valley. (jûluts dal.) junction peak 71ø (71°51.1´n 25°13.2´w; map 5). peak in the upper reaches of roslin gletscher, stauning alper. named by the 1970 cambridge university expedition which climbed the mountain on 1 august 1970. junctiondal 73ø-559 (73°15.3´n 25°54.7´w; map 4). valley in southern andrée land, named by j.m. wordie’s 1929 expedition as junction valley, because the valley followed fault contacts between different rock units. on some maps a hunting hut is shown at the mouth of the valley, but this was never built; the material left here was removed in 1935 to build a hut in nordfjord (p.s. mikkelsen 1994). jupiter gletscher 71ø-331 (71°42.3´n 25°10.8´w; map 5). glacier fig. 49. the 1300 m high south face of the mountain jomsborg, on the west side of renbugten. the lower half of the cliff comprises light coloured foliated granite. 210 flowing ne to join bjørnbo gletscher, south stauning alper. named jupiter glacier by john hunt’s 1960 expedition, for the planet jupiter, fifth major planet from the sun. juradal 71ø (71°19.6´n 22°38.6´w) name used on the maps of callomon (1970) for the valley in ne jameson land carrying liaselv, which flows east into carlsberg fjord. the valley was used during lauge koch’s 1958 expedition as a route to the interior of jameson land, and named after the jurassic age of the rocks. juraelv 71ø-191 (72°06´n 24°04´w; map 4). river in west jameson land draining sw into lodin elv. named by hans stauber during lauge koch’s 1936–38 expeditions after the widespread outcrops of jurassic rocks. juraklöft 74ø (74°39´n 20°15´w). name used by maync (1947) for a ravine in north wollaston forland, just east of sillerendal. the name arose during lauge koch’s 1936–38 expeditions, and was given for the jurassic rocks. jyllandselv 70ø-94 (70°46.1´n 23°41.1´w; map 4). river in sw jameson land flowing sw to enter the sea north of vandre blokken. named during the 1931–34 treårsekspeditionen by lau rits bruhn for jylland (= jutland), denmark. jytte havn 71ø-417 (71°03.5´n 25°37.4´w; fig. 50). pronounced bay in the sw island of the bjørneøer, regularly used as an anchorage by ggu’s 9-ton motor cutter jytte during the 1967–72 ggu scoresby sund expeditions. jægerdal 72ø-398 (72°06.5´n 23°37.6´w). valley in north scoresby land, sw of mesters vig. named by hans kapp during lauge koch’s 1957–58 expeditions, and named after the norwe gian hunters (= jæger). jægerly 74ø (74°59.0´n 18°23.7´w). name sometimes used for the danish hunting hut built in 1923 by østgrønlands fangst kom pagni about 2 km ne of kap david gray, shannon (see kap david grayhytten). jægersund 76ø-172 (76°18.8´n 20°40.0´w; map 4). sound between tvillingerne and nanok ø, in the sw part of dove bugt. the name was suggested by the place name committee as a substitute for a proposal by eigil knuth. it commemorates the danish hunters (= jæger) operating in the region. a danish hut on the south point of nanok ø, known officially as hasserishytten, is also known as jæger sundhytte or sydlige jægersundhytte. engelhards sund has also been used. jägmästeren ø 72ø-32 (72°28.4´n 24°41.8´w; map 5). island at the mouth of segelsällskapet fjord. the name jägmästerens ö was originally given by a.g. nathorst’s 1899 expedition to the present karlenes ø, and commemorated e. nilson [b. 1863], the expedition hunter, always referred to in the expedition narrative as ‘jäg mästeren’. koch (1929a) extended the original usage to four large islands and several small skerries which he termed jagmasters islands. in time the name became attached to the present island. (forest officer island, jägmästernsö.) jættebringen 77ø-39 (77°23.7´n 23°50.6´w; map 4). eastern part of ymer nunatak in north dronning louise land. so named by 1906–08 danmark expedition probably for its shape, and perhaps also after the feature of the same name at møns klint, denmark. koch (1916) translates it as ‘the giant’s chest’. (jættebrinken.) jættedal 73ø (73°28.7´n 25°58.5´w). major valley in louise boyd land draining east to jættegletscher. the name was used by jan escher describing geological fieldwork in 1997–98. jættedal 70ø-192 (70°31.6´n 22°05.1´w). major valley in southern liverpool land draining into hvalrosbugt, so named during the 1931–34 treårsekspeditionen by laurits bruhn for its size (jætte = giant). a rough landing strip here, suitable for small aircraft, was used for many years as a means of access to nearby scoresbysund. jætteelv 70ø-191 (70°31.6´n 22°05.1´w). river draining through jættedal in southern liverpool land. so named by laurits bruhn during the 1931–34 treårsekspeditionen. jættegletscher 73ø-521 (73°27.0´n 27°37.0´w; map 4; see also fig. 58). glacier at the head of isfjord, between frænkel land and louise boyd land. named jätteglacieren by a.g. nathorst’s 1899 expedition because it gave rise to the very large icebergs in isfjord (jætte = giant). (jatte glacier, giant glacier, jettebreen, jaette gletscher.) jættehorn 73ø-676 (73°33.0´n 26°08.6´w). mountain in central andrée land, on the south side of grejsdal. named by john haller following explorations during lauge koch’s 1949–51 expeditions, because of its large size and central spire. jættevæggen 75ø-77 (75°11.0´n 22°27.7´w; map 4). impressive cliff on the north side of heinkel gletscher and inner grandjean fjord. the name originated from the wintering party at kulhus during the 1931–34 treårsekspeditionen. unpublished maps show the original field name to have been the seven pillars of hell. jøkelbugten 78ø-11 (78°25.0´n 20°20.0´w; maps 1, 4). extensive bay east of hertugen af orléans land. the name was originally fig. 50. ggu’s small cutter jytte painted in the traditional deep red colour sailing with geological parties in the east greenland fjords. 211 used by the 1906–08 danmark-ekspeditionen in the form jøkel bugt and jökelbugten, and arose because the bay was covered by an essentially connected mass of floating glacier ice, extending out to the outlying row of islands and skerries. ‘jökel’ is old norse for a glacier, a form still in use in iceland. (jökel bay, jøkel bay.) jötulen 71ø (71°36.8´n 23°13.5´w). mountain sw of the head of fleming fjord between rhætelv and enhjørning dal. named by the norwegian hunters helge ingstad and normann andersen during their 1932–34 expedition because of its ominous appearance and curious reddish colour (jötulen = ogre). (ogre mountain.) k kaasarip nasaa [storø] 70ø-6 (70°49.5´n 27°30.0´w). large is land on the east side of rødefjord. one of the names recorded by the 1955 geodætisk institut name registration, it translates as ‘kejsers (emperor’s) hat’, and derives presumably from the shape. (kaisarip naâ.) kai nielsen fjeld 79ø-16 (79°25.6´n 20°41.6´w; map 4). mountain in northern lambert land. named by the 1938–39 mørkefjord expedition after the danish sculptor kai nielsen [1882–1924], and described as a “steep mountain on the north side of lamberts land” (knuth 1942). it was probably intended for the 1023 m mountain known as trompeteren bastion, but on official maps it has been misplaced westwards to the not very conspicuous north ‘cape’ of lambert land and with the erroneous spelling kap nielsen fjeld. kai nielsen’s works include the monument to mylius-erichsen, høeg-hagen and jørgen brønlund of the 1906– 08 danmark-ekspeditionen, erected on langelinje, copenhagen. kaisarip nasâ – see kaasarip nasaa. kajkap 77ø-56 (77°19.3´n 18°55.0´w). cape on the south side of skærfjorden. so named by david malmquist during the 1931–34 treårsekspeditionen after karin (kaj) lunell (born hadders), his wife’s sister. kaka 72ø (72°42.1´n 22°50.9´w). small island in vega sund, south of the scott keltie øer, the present thora ø. so named on the nsiu maps of lacmann (1937), for the shape (kaka = cake). kaldbakur 76ø-139 (76°31.4´n 26°10.5´w; map 4). nunatak in sw dronning louise land, named by j.p. koch’s 1912–13 expedition. the expedition was delayed here for seven days by bad weather, and koch (1913) records the name is icelandic for ‘cold hill’. kalebarodden – see kobberpynt. kalifbjerg 73ø-708 (73°09.5´n 28°40.4´w; map 4). mountain 2667 m high in the nunatak region of western frænkel land near peter mann bjerg. named by john haller and eduard wenk during explorations on lauge koch’s 1951 expedition, because the windpacked snow collapsed under their weight such that the only way of progressing was upon their knees, as if approaching a caliph. (kalif bjerg ). kalkdal 70ø-160 (70°50.2´n 22°15.0´w; map 4). valley in liver pool land east of fame øer. named in the geological account of g.c. amdrup’s 1898–1900 expedition as kalkdalen or limestone valley, for the occurrence of limestone. the name was not used on maps until 1934 when it was revived and approved at the suggestion of brian roberts. it is used as a sledge route between hurry inlet and the eastern outer coast of liverpool land. kalkdalen 70ø (70°47.8´n 22°26.3´w). a hut built by scoresbysund municipality south of the mouth of kalkdal is known as kalkdalen or gåsereden (goose’s nest), and by the greenlandic name kanger saaiva. sandell & sandell (1991 p. 96) use the name nerterit inaat for this hut, which they report has been used for char-fishing. kalles hytte 74ø (74°01.4´n 22°17.8´w). norwegian hunting hut on the south side of wordie bugt, 2 km west of surprise elv. erected by finn devold’s expedition in 1929, and named after karl nico laisen who helped build it. it is also known as wordie bugt hytten. kalotten 74ø-299 (74°47.0´n 20°56.9´w). mountain about 1000 m high in th. thomsen land. the name originated from the wintering parties at eskimonæs and kulhus during the 1931–34 treårs eks peditionen, and was given because the ice cap on the summit resembles a skull cap (= kalot). kalsneset 72ø (72°41.2´n 22°12.5´w; fig. 14). cape on se geo graph ical society ø on the north side of vega sund. so named on nsiu maps of lacmann (1937), for the locality of the same name in vesterålen, norway. kalvedal 73ø-643 (73°32.1´n 26°44.8´w; map 4). valley in sw andrée land draining south to rendal. named during the 1931– 34 treårsekspeditionen by ove simonsen because many new-born musk-ox calves were seen here. kalven 74ø-266 (74°00.6´n 20°56.3´w; map 4). island in the finsch øer group, south of store finsch. the name first appears on an nsiu map (1932a), and derives from its small size relative to store finsch (kalven = the calf ). kalven 76ø-52 (76°55.2´n 20°33.1´w). island in mørke fjords bug ten. so named by the 1906–08 danmark-ekspeditionen (kalven = the calf ). (kalvenø, kálfur.) kalvodden 76ø-269 (76°55.4´n 20°39.2´w). headland on the north side of vædderen, opposite the west end of kalven. named by the 1938–39 mørkefjord expedition. kamelen 73ø-254 (73°00.9´n 22°11.7´w). small island in the broch øer group. the name appears on an nsiu map (1932a), and de rives from a resemblance to the hump of a camel. kamelen 74ø (74°29.4´n 20°31.8´w). hill with a pronounced hump, nne of zackenberg forskningsstation kamelgletscher 70ø-346 (70°05.4´n 22°28.0´w). glacier on the ne side of kamelryggen, savoia halvø. named during the 1931– 34 treårsekspeditionen by laurits bruhn. kamelryggen 70ø-345 (70°05.1´n 22°28.8´w). mountain ridge with two summits (or humps) 1037 m and 900 m high on savoia halvø. named during the 1931–34 treårsekspeditionen by laurits bruhn for its shape. kamelryggen 74ø (74°10.4´n 20°13.8´w). hill with two summits on east clavering ø, probably the present magnetikerbjerg. the name came into use among hunters of østgrønlandske fangstkompagni about 1923 due to its shape in profile. (camel hill.) kamæleon 76ø-329 (76°29.5´n 25°55.3´w; map 4). nunatak in sw dronning louise land between kaldbakur and gefiontinder. named by the 1952–54 british north greenland expedition be cause of its variable colour in different lights, after the chameleon. from the north it appears to be a pyramid, but is in fact a long n–strending ridge. kanaans land 71ø (71°18.0´n 24°00.0´w). name used by ingstad (1935) for the area of jameson land sw of olympen, which helge ingstad and normann andersen had been unable to reach in the spring of 1932 because of deep snow. they had to cross this region to reach the interior of nordvestfjord, the ‘promised land’ lying north of 71°30´n latitude and thus within the boundaries of their eirik raudes land. kangerdlugssuaq – see kangerlussuaq. kangerlussuaq / kangertittivaq [scoresby sund] 70ø-258 (70°17.0´n 23°00.0´w). very large e–w-trending fjord. recorded by the 1955 geodætisk institut name registration, the name translates as ‘the big fjord’. the west greenland form kangerdlugssuaq (= kangerlussaq) was said to be in use in 1955, whereas kanger títivak (now kangertittivaq) was said to be the official name. both these names have approved status. some recent maps (e.g. tuborg & sandell 1999) give the spelling kangersuttuaq. kangersaaiva – see kalkdalen. kangersik kiattek [nordvestfjord] 71ø-37 (71°31.0´n 26°00.0´w; map 4). name found on modern maps, that replaces kanger tertivarmiit kangertivat recorded by the 1955 geodætisk institut name registration. kangerstua 71ø (c. 71°17´n 25°04´w). name occasionally used for 212 the larger of the two houses built at sydkap. see also kanger tertivarmiit. kangersuttuarmiit kangersuat – see kangertertivarmiit kangertivat. kangersuttuaq – see kangerlussuaq. kangersuttuup kangersua – see kangertertivarmiit kangertivat. kangerterajígtap igtiva – see kangerterajittap ittiva. kangerterajigtap ilivnera – see kangerterajittap ilinnera. kangerterajitta itterterilaq [carlsberg fjord] 71ø-46 (71°25.6´n 22°24.1´w). fjord between jameson land and northern liverpool land. the name was recorded by the 1955 geodætisk institut name registration, and translates roughly as ‘our little fjord’s equivalent on the other side’. this refers to the relative positions of carlsberg fjord and hurry inlet north and south of kanger terajittap ilinnera [klitdal]. (kangerterajivta igterterilâ.) kangerterajittap ilinnera [klitdal] 70ø-118 71ø-125 (70°59.4´n 22°29.4´w). low valley providing an easy sledge route north from the head of kangerterterajiva [hurry inlet]. one of the names recorded by the 1955 geodætisk institut name registration, it translates roughly as ‘the crossing place at the head of the little fjord’. (kangerterajigtap ilivnera.) kangerterajittap ittiva 70ø-155 (c. 70°51´n 22°28´w). hunting hut at ulveodde, at the head of hurry inlet. the name was recorded by the 1955 geodætisk institut name registration, and can be translated as ‘our hut in kangerterajiva’. it was said in 1971 to have disappeared. (kangerterajígtap igtiva.) kangerterajiva [hurry inlet] 70ø-148 (70°36.0´n 22°31.0´w). fjord between south liverpool land and jameson land. one of the names recorded by the 1955 geodætisk institut name registration, it translates as ‘the little fjord’. kangerterajiva [nordostbugt] 71ø-33 (71°20.0´n 24°41.0´w; map 4). bay or short fjord east of sydkap at the mouth of schuchert flod. recorded by the 1955 geodætisk institut name registration, the name means ‘the little bay’. the greenlandic name marrakajik (= ‘the little clay’) has also been used for the bay, but is correctly the approved name for the nearby large river (schuchert flod). kangerterajiva 71ø-208 (71°18.4´n 25°05.6´w). bay west of syd kap at the mouth of nordvestfjord. recorded by the 1955 geo dætisk institut name registration, the name means ‘the little bay’. kangerterajivta igterterilâ – see kangerterajitta itterterilaq. kangertertivarmît – see kangertertivarmiit. kangertertivarmît kangertivat – see kangertertivarmiit kangertivat. kangertertivarmiit [sydkap] 71ø-212 (71°17.3´n 25°04.5´w). inuit (eskimo) settlement at sydkap, at the mouth of nordvestfjord. there are many ruins in the vicinity and on outlying islands, and the locality was periodically occupied by greenlanders from scores bysund between 1934 and 1954. two stone built houses were built here in 1946, one a shop and store house. recorded by the 1955 geodætisk institut name registration, the name translates as ‘those that live at the big fjord’. the name has also been reported as kangerstua, a reference to the larger of the houses. (kanger terti varmît.) kangertertivarmiit kangertivat [nordvestfjord] 71ø-37 (71°31.0´n 26°00.0´w). large and very long fjord arm. this name was recorded by the 1955 geodætisk institut name registration for nord vestfjord, and the name means ‘kangertertivarmiit’s big fjord’. tuborg & sandell (1999) record the variations kangersuttuarmiit kangersuat and kangersuttuup kangersua, and on modern maps the name kangersik kiattek is used. (kangertertivarmît kangertivat.) kangertítivak – see kangertittivaq. kangertittivaq / kangerlussuaq [scoresby sund] 70ø-258 (70°17.0´n 23°00.0´w). very wide, e–w-trending fjord. one of the names recorded by the 1955 geodætisk institut name registration, the name means ‘the big fjord’. in 1955 kangertítivak (now kangertittivaq) was said to be the official name, although the west greenland form kangerdlugssuag (now kangerlussuaq) was also in use. current local usage is said to be kangertittivaq, and both names have approved status. the variation kangersuttuaq was recorded by tuborg & sandell (1999). kangertivatsiaakajiip nuaa 70ø-211 (70°36.8´n 21°42.1´w). cape on the south side of kangertivatsiaakajik [lillefjord]. one of the names recorded by the 1955 geodætisk institut name registration, it translates as ‘the large bay’s cape’. (kangertivatsiâkajîp nûa). kangertivatsiaakajik [hartz vig] 70ø-332 (70°26.6´n 21°48.8´w). bay between kap tobin and kap swainson. recorded by the 1955 geodætisk institut name registration, the name translates as ‘the rather large bay’. the local scoresbysund newspaper recorded in 1984 the name nappangulikajiip kangersiva for this bay. (kanger tivatsiâkajik). kangertivatsiaakajik [lillefjord] 70ø-212 (70°37.8´n 21°40.7´w). large bay or small fjord in se liverpool land. recorded by the 1955 geodætisk institut name registration, the name means ‘the little fjord’. (kangertivatsiâkajik). kangertivatsiâkajîp nûa – see kangertivatsiaakajiip nuaa. kangertivatsiâkajik – see kangertivatsiaakajik. kangertivit anginersaat [storefjord] 71ø-129 (71°05.4´n 21°54.6´w). largest of the fjords on the east coast of liverpool land. recorded by the 1955 geodætisk institut name registration, the name means ‘the largest of fjords’. (kangertivit anginersât). kangertivit anginersât – see kangertivit anginersaat. kangigai – see kangikajik [kap brewster]. kangikajiip appalia [kap brewster] 70ø-361 (70°08.8´n 22°04.5´w). prominent headland on the south side of scoresby sund, whose alternative approved name is kap brewster. until 1978 the authorised greenlandic name was kangikajik, but al though this name is still found on some modern maps (e.g. tuborg & sandell 1999), it is now officially applied to the settlement west of the cape. the present name translates as ‘little auk’s cape’, and refers to the bird colonies on the cliffs. kangikajiip kangerterajiva [vikingebugt] 70ø-75 (70°19.1´n 25°14.2´w). bay or short fjord east of kap stevenson (kangikajik) on the south coast of scoresby sund. one of the names recorded by the 1955 geodætisk institut name registration, it translates as ‘kangikajik’s bad fjord’. (kangikajîp kangerterajiva.) kangikajik [kap brewster] 70ø-355 (70°07.8´n 22°14.5´w). settlement west of kap brewster, occupied periodically. until 1978 the authorised name of the settlement was kangikajingmît, the present name kangikajik formerly being applied to the cape itself (see kangikajiip appalia). kangikajik translates roughly as ‘the bad cape’. on some recent maps (e.g. tuborg & sandell 1999) the name kangikajik is still used as the greenlandic name for the cape. recent reports suggest there is only one habitable house, that goes by the name kangigai. (kangikaiimit.) kangikajik [kap stevenson] 70ø-73 (70°24.4´n 25°12.3´w). prominent headland on the south coast of scoresby sund. the name was recorded by the 1955 geodætisk institut name registration, and translates as ‘the bad cape’. some modern maps record kangikajik kangitteq (tuborg & sandell 1999). kangikajingmît 70ø (70°08´n 22°16´w). this was formerly the authorised name for the settlement west of kap brewster, and was that recorded by the 1955 geodætisk institut name registration. it translates as ‘those that live at the bad cape’. in 1978 the authorised form was changed to kangikajik to comply with current usage by the inhabitants. recent reports suggest there is only one habitable house, that goes by the name kangigai. kangikajîp agpalia – see kangikajiip appalia. kangikajîp kangerterajiva – see kangikajiip kangerterajiva. kangoq ryg 81ø (81°09.2´n 13°18.4´w). ridge of moraine in eastern kilen, kronprins christian land. the name is found on a coloured geological map of kilen printed in 1991 (pedersen 1991), and is the greenlandic name for the barnacle goose. kap – see also cap (old danish, german), cape (english), kapp (nor wegian). 213 kap achton friis 76ø-121 (76°46.2´n 23°04.6´w; map 4). capelike feature in eastern dronning louise land, named by j.p. koch’s 1912–13 expedition after [johannes] achton friis [1871–1929], danish artist on the 1906–08 danmark-ekspeditionen. together with aage bertelsen he made several hundred paintings and drawings during the expedition. some of friis’ portraits are included in koch (1913). friis was also an author and wrote a popular book on the 1906–08 danmark-ekspeditionen, but was most noted for his books on denmark, particularly ‘de danskes øer’ published in 1926. kap adam 74ø-273 (74°11.4´n 22°13.1´w). cape on the north side of hansen havn, north of jordan hill. named, probably by norwegian hunters, as kapp adam, the name first appearing on an nsiu map (1932a) together with kapp eva, now kap eva. kap adolf jensen 79ø-5 (79°41.4´n 20°00.0´w; map 1, 4). sw cape of hovgaard ø. so named by the 1906–08 danmark-ekspe ditionen for adolf severin jensen [1866–1953], a danish zoologist who assisted mylius-erichsen in the planning of the 1906–08 danmark-ekspeditionen. see also ad. s. jensen land. kap agnes – see agnes-tufta. kap ahrens 76ø (c. 75°53´n 19°47´w). name given to a cape on the east coast of hochstetter forland north of haystack by karl koldewey’s 1869–70 expedition. the name occurs only in the article by lenz (1874) (j. løve, personal communication 2010). kap alf trolle 75ø-90 (75°55.9´n 18°28.7´w; map 2, 4). southernmost low cape of store koldewey, south of kap arendts (it has also been called kap arendts næs). the name was given by the 1932 gefion expedition after alf trolle [1879–1949], captain of the expedition ship during the 1906–08 danmark-ekspe ditionen. he was also deputy leader and hydrographer, and had made astronomical observations at the cape during the expedition. later trolle also took part in the 1932 gefion expedition. in 1933 he founded with his wife a memorial fund, which published a series of reports relevant to the 1906–08 danmark-ekspeditionen and hunting and exploration in east greenland. kap alfred 72ø-42 (72°49.9´n 25°33.2´w; map 4; fig. 29). north ern cape of lyell land at the mouth of kempe fjord, named by a.g. nathorst’s 1899 expedition. the name was probably given for a member of nathorst’s family, as were many other capes in the vicinity. (cape alfred.) kap alfred beauvais 76ø-84 (76°42.0´n 18°43.9´w). cape on the east side of lille koldewey. so named by the 1906–08 danmarkekspeditionen for the supplier of provisions to the expedition. alfred beauvais was a director of the meat-packing company danica which supplied this and many other danish expeditions (j. løve, personal communication 2009). kap allen 71ø-14 (71°41.0´n 22°00.0´w; map 4). cape on canning land named by william scoresby jr. in 1822 as cape allan for an edinburgh friend. this was probably thomas allan [1777–1833], a mineralogist who had purchased giesecké’s greenland mineral collection, which among other items included a new mineral called after its purchaser, allanite. scoresby used the spelling ‘allen’ on his chart in error, and it is this form that is invariably used on maps today. the german edition of scoresby’s narrative (1825) uses the correct ‘allan’ form throughout. (cape allen, cape allén.) kap amélie 77ø-8 (77°31.1´n 19°13.0´w; map 1, 2, 4). cape north of the mouth of penthievre fjord. named by the duke of orléans in 1905 as cap amélie, possibly for marie amélie des deux siciles [1782–1866], wife of louis-philippe 1, king of france. (cap amelie.) kap amélie hytten 77ø (77°32.1´n 19°08.0´w). hut built on 15 march 1941 about 3 km ne of kap amélie by ib poulsen and other meteorologists of eigil knuth’s 1938–39 expedition. the outbreak of war in europe had disrupted normal contacts with denmark, but meteorological observations were continued at mørkefjord station. the hut was still erect in 1990, missing only the lower parts of the north and south walls. a newer sirius hut is found at the same locality. kap amundsen 78ø-43 (78°56.6´n 18°03.3´w; map 1, 4). south cape of the southernmost island in the norske øer. named by john haller following explorations during lauge koch’s 1956–58 expeditions after roald amundsen [1872–1928]. a noted norwegian polar explorer, amundsen’s first major exploit was the traverse of the nw passage in the gjöa in 1906–08, followed up by the first attainment of the south pole in june 1910. he also took part in the first flight over the north pole in 1926 with lincoln ellsworth and umberto nobile, and disappeared in 1928 during an attempt to rescue nobile whose airship had crashed off west spitsbergen. kap anatektite – see kap jones. kap anna bistrup 79ø-7 (79°41.0´n 18°14.1´w; map 1, 4). se cape of hovgaard ø. so named by henning bistrup, during the 1906–08 danmark-ekspeditionen after his mother anna vilhelmine augusta østerberg [1848–1934]. see also l. bistrup bræ. kap arendts 75ø-88 (76°05.9´n 18°35.7´w; map 4). name used for the mountain north of kap alf trolle on store koldewey. it was named by karl koldewey’s 1869–70 expedition as cap arendts for karl arendts [1815–81], german geographer, professor at munich, and amongst the founders of the geographical society of munich. he was a supporter of the expedition and had formed a committee to raise funds. kap arendts næs was used by the 1906–08 danmarkekspeditionen for the low peninsula between kap arendts and kap alf trolle. kap arnakke 74ø-107 (74°11.3´n 20°07.1´w). east cape of clavering ø. the name first appears on a sketch map in gustav thostrup’s 1921 logbook (møller 1939) in the form c. arnak. (cap arnak, cape arnak.) kap barclay 69ø-11 (69°16.5´n 24°36.0´w). cape on the northern blosseville kyst, named by william scoresby jr. in 1822 as cape barclay, after john barclay [1758–1826]. he was a noted anato mist, and lecturer at the edinburgh college of surgeons from 1804. kap basel 71ø (71°51.5´n 28°54.7´w). name used for the northern peninsula of hinks land, in the report on the 1931–34 treårsekspeditionen by helge g. backlund (in: koch 1955). the location is referred to by haller (1971). named after the swiss city of basel, the home town of eduard wenk, a member of backlund’s 1933 party. kap bayard 72ø-409 (72°46.4´n 26°25.2´w; map 2; see also fig. 52). cape between the mouths of dickson fjord and röhss fjord. named by a.g. nathorst during his 1899 expedition, probably for a.g. bayard, a stockholm engineer who had contributed 400 swedish kronor to the expedition finances. kap beijer 72ø-410 (72°46.6´n 26°17.6´w; map 4; see also fig. 52). cape in south suess land, east of the mouth of dickson fjord. named by a.g. nathorst during his 1899 expedition for gottried beijer [1838–1901], a successful malmö businessman who contributed 600 swedish kronor to the expedition finances. beijer was noted as one of the founders of modern malmö. (cape beijer.) kap bellevue 77ø-38 (77°05.2´n 23°12.2´w; map 4). cape-like prominence or mountain in dronning louise land on the west side of storstrømmen. so named by the 1906–08 danmark-ekspedi tionen for the spectacular view. kap bennet 73ø-10 (73°23.4´n 21°35.5´w; maps 3, 4). cape on the south side of mackenzie bugt. william scoresby jr. named a feature in this area as bennet island, after captain bennet of the venerable, one of the group of whalers that accompanied scoresby in 1822. the location of scoresby’s original ‘island’ is uncertain, although he placed it north of the mackenzie bay of his chart. the name was transferred to a cape at the present location on koldewey’s (1874) maps, and although moved north of mackenzie bugt on subsequent danish maps, it was later moved back to the present site. scoresby probably could not have seen the present kap bennet. norwegian hunters have occasionally called 214 the cape giskeodde. (cape bennet, cape bennett, halbinsel bennet, kapp bennet, bennet hill.) kap bennet hytte – see bennethytta. kap bergendahl 78ø-6 (78°37.7´n 18°22.3´w; map 4). east cape of one of the franske øer. named by the duke of orléans in 1905 as cap bergendahl for r.-svante bergendahl, a lieutenant in the swedish navy who was one of the officers on the expedition ship. kap bergendahl 79ø (79°09.3´n 19°04.0´w). name used for the east cape of lambert land in the popular account of the 1906–08 danmark-ekspeditionen by friis (1909). as this name had previously been used for a cape in the franske øer (see above) it was discontinued for this site, which is now known as brønlunds grav). kap berghaus 74ø-61 (74°16.8´n 20°09.0´w). cape in sw wolla ston forland. named by karl koldewey’s 1869–70 expedition as cap berghaus, perhaps after hermann berghaus [1828–90], a ger man cartographer at justus perthes geographisches anstalt in gotha, publishers of petermanns mitteilungen. possibly also named after heinrich berghaus [1797–1884], one of the initiators of the berlin geographical society. norwegian hunters have used heklas hvalrossnæss for the same feature. (cape berghaus). kap berghaushytten 74ø (74°16.9´n 20°07.8´w). danish hunting hut immediately east of kap berghaus, south wollaston forland, built by nanok in september 1946. (kap berghaus hytten.) kap bergliot 77ø-31 (77°30.0´n 20°09.1´w). cape between assut sund and h.g. backlund fjord, west of skærfjorden. named by alf trolle during the 1906–08 danmark-ekspeditionen for his future wife bergliot holm [1885–1943]. kap berlin 74ø-37 (74°41.0´n 19°25.7´w; maps 2, 4). cape in northern wollaston forland. named by karl koldewey’s 1869–70 expedition as cap berlin, after the capital city of the north german federation. a large collection of donations in support of the expedition was made in berlin. (cape berlin.) kap berlin hytte – see berlin-stua. kap bernhoft 79ø-11 (79°42.0´n 20°39.0´w). cape in southern kronprins christian land, nw of nioghalvfjerdsfjorden. mapped from the air by lauge koch during the 1931–34 treårsekspedi tionen, it was named after h.a. bernhoft [1869–1958]. bernhoft was director of the danish foreign ministry during the dispute over east greenland decided at the international court at the hague. kap beurmann 76ø-1 (76°03.2´n 19°47.8´w; map 4). cape on the north side of the mouth of bessel fjord. named by karl koldewey’s 1869–70 expedition as cap beurmann, for karl moritz von beur mann [1835–63], a german african explorer. the norwegian 1932–34 giæver expedition built a hut in the bay west of the cape. (cape beurmann point, kap beurmanns næs.) kap biot 71ø-21 (71°54.5´n 22°31.8´w; map 4). cape north of fleming fjord. it was named by william scoresby jr. in 1822 as cape biot, in compliment to the french philosopher and astron omer, jean baptiste biot [1774–1862]. kap biot 71ø (71°52.8´n 22°39.2´w). danish hunting station built in september 1940 on the nw side of fleming fjord at the foot of kap biot. the station and personnel were brought up by the furenak with the intention of establishing a weather station for support of german activities in the north atlantic. the norwe gian ship fridtjof nansen, in the service of the united states, evacuated the personnel and burnt the station on 7 september 1940. kap bismarck 76ø-10 (76°42.0´n 18°33.0´w; map 4). southern most peninsula of germania land. named cap bismarck by karl koldewey’s 1869–70 expedition, after otto eduard leopold von bismarck [1815–1898]. bismarck was at the time chancellor of the north german federation, and was present with kong wilhelm when the expedition sailed from bremerhaven on 15 june 1869. koldewey’s original cap bismarck was said to be the south spur of harefjeldet according to koch (1916 p. 374). the name was used for the low cape at the present position by the duke of orléans in 1905, a position retained by the 1906–08 danmark-ekspeditionen. (cape bismarck mountain, bismarckshöfdi.) kap bismarck hytten 76ø (76°42.0´n 18°33.0´w). hut at kap bismarck built for danmarkshavn weather station in 1979. kap bismarck-næsset 76ø (76°42.9´n 18°33.6´w). name used by friis (1909) in his popular account of the 1906–08 danmarkekspeditionen for the low peninsula of which the present kap bismarck forms the south end. (kap-bismarck-tangen.) kap bjarne nielsen 76ø (76°36.9´n 21°00.4´w). prominent ne cape of edvard ø, in dove bugt. so named by the 1932 gefion expedition after bjarne nielsen [1876–1953], a businessman and general consul, who was a member of the board of østgrønlandske fangst kompagni nanok 1929–36. the name is used in den grønlandske lods (1968). kap borlase warren 74ø-20 (74°16.0´n 19°22.7´w; map 4). cape in se wollaston forland, named cape borlase warren by douglas clavering in 1823. it was the first place at which clavering landed, and was named after sir john borlase warren [1753–1822], who in 1780 had married caroline, a daughter of sir john clavering. the norwegian and danish hunting huts built at the cape have been known by various names: kap borlase warren hytten, kap borlase warren station, borganes, valdermarshaab, grønlænderhuset, sverd rups nes, bjørn-heimen. (c. borlace warren, cape borlase warren.) kap bornholm 76ø-85 (76°43.8´n 18°49.0´w). northern cape of lille koldewey. so named by christian b. thostrup during the 1906–08 danmark-ekspeditionen, after herman koefoed, a mem ber of the expedition who was educated and worked on the danish island of bornholm, and was thostrup’s faithful assistant during preparation of detailed maps (thostrup 2007). kap bourbon 78ø-5 (78°44.7´n 18°06.4´w; map 4). east cape of bourbon ø, the northernmost point seen by the duke of orléans in 1905, who named it cap bourbon. the orléans family was linked to the bourbons through louise-philippe 1 of france, great-grandfather of the duke of orléans. kap bratthuken – see bratthuken. kap bremen 74ø-24 (74°58.9´n 19°58.3´w; maps 2, 4). cape on ne kuhn ø. named by karl koldewey’s 1869–70 expedition as cap bremen, for the city of that name. bremen was the home of the principal supporting committee of the expedition, ‘der verein für die deutsche nordpolarfahrt in bremen’, which was responsible for publication of koldewey’s narrative. the senate of bremen made substantial donations to the expedition finances. kap bremen hytten 74ø (74°59.0´n 19°58.2´w). danish hunting hut on the south side of kap bremen, ne kuhn ø, built by nanok in september 1931. now a ruin. (kap bremenhytten.) kap breusing 74ø-75 (74°12.7´n 20°06.8´w). cape on east clavering ø. named by karl koldewey’s 1869–70 expedition, and on different maps spelt cap breusing or cap breussing. the name was evidently intended to honour friedrich august arthur breu sing [1818–92]. a german nautical expert, and director of the naval academy at bremen from 1858, he played a leading part in the organisation of german polar expeditions. cap holcha has been used for the same feature by danish hunters. (kapp breusing.) kap breusinghytten 74ø (74°12.6´n 20°07.0´w). hut on the south side of kap breusing, built in 1951 by daneborg weather station personnel using material from the wartime american station at dødemandsbugten. kap brewster [kagikajiip appalia] 70ø-361 (70°09.0´n 22°03.5´w; maps 3, 4). prominent cape on the south side of the mouth of scoresby sund. it was named cape brewster by william scoresby jr. in 1822 in compliment to a much esteemed friend, david brewster [1781–1868]. brewster was very active in scientific circles, pub lished many papers on the polarisation of light, and invented the kaleidoscope. kap brewster [kangikajik] 70ø-355 (70°07.8´n 22°14.5´w). 215 settlement west of kap brewster periodically occupied by families from scoresbysund. until 1978 the authorised name was kangi kajingmît, the present name kangikajik formerly being applied to the cape itself (see kangikajiip appalia). kangikajik translates roughly as ‘the bad cape’. (kangikaiimit.) kap broer ruys 73ø-3 (73°31.8´n 20°22.8´w; maps 2–4). cape in se hold with hope. the name was applied to the cape by douglas clavering, who climbed to an adjacent summit (rochusspids) on 7 september 1823. this cape approximately corresponded with the position of cape broer ruys of old dutch charts, a name originating with a sighting by a whaler in 1655. the name appears on charts in the collection by van keulen dating from 1681 and 1706 as t'land v. broer ruys. cape hold with hope was used for the same general fea ture by whalers in the 19th century, and norwegian hunters in the 1930s. the grave of a norwegian hunter, john tutein, killed by a bear in 1921 is found here. kap broer ruys nord – see broer ruys nord. kap broer ruys station 73ø (73°29.0´n 20°25.3´w). danish hunting station in sw hold with hope. built by østgrønlandske fangst kompagni in 1920 and taken over by nanok in 1929, it was severely damaged by a storm in the winter of 1936–37. it has also been known as cape hold with hope, kap hold with hope station and station ‘b’. kap broer ruys syd – see broer ruys syd. kap brown 71ø-17 (71°47.5´n 22°25.6´w; map 4). cape at the north point of wegener halvø. it was named cape brown by william scoresby jr. in 1822 after the botanist robert brown [1773–1858], who published memoirs on australian plants, and became keeper of botanical collections at the british museum. kap brown hytten – see brown-stua. kap brown huset 71ø (71°43.2´n 22°43.9´w). small wintering station on the east side of fleming fjord, 15 km sw of kap brown, built by the 1931–34 treårsekspeditionen in august1931. it has also been known as vimmelskaftet and flemmingfjordhuset. kap buch 75ø-19 (75°08.3´n 20°30.5´w; map 4). cape between the mouths of ardencaple fjord and grandjean fjord. named cap buch by karl koldewey’s 1869–70 expedition, after the noted german geologist and palaeontologist christian leopold von buch [1774–1853]. von buch was responsible for the first geological map of germany, and played an important part in the controversies of wernerism and neptunism. see also buch bjerg. kap buchenau 74ø-38 (74°43.4´n 18°34.0´w; map 4). northern cape of lille pendulum. named cap buchenau by karl koldewey’s 1869–70 expedition, for franz georg philipp buchenau [1831– 1906]. a botanist, he was professor of a school in bremen, and had prepared one of the botanical sections of koldewey’s narrative (j. løve, personal communication 2010). (kap buchi now.) kap buchenau hytten 74ø (74°43.4´n 18°33.6´w). danish hunting hut at kap buchenau, nw lille pendulum. built by nanok in 1930. (kap buchenauhytten.) kap buchhytten 75ø-104 (75°11.4´n 20°34.4´w). danish hunting hut about 6 km north of kap buch, on the south side of ardencaple fjord, built by nanok in 1932. only the stone foundations of the hut remain (p.s. mikkelsen 2008). (kap buch hytten, kap buch hytta.) kap buddicom 71ø-128 (71°04.5´n 21°41.4´w; map 4). cape on the east coast of liverpool land south of storefjord, named cape buddicom by william scoresby jr. in 1822 in compliment to a respected clergyman of liverpool. scoresby took lessons in latin and greek from revd buddicom in 1823, the first essentials towards ordination. kranck (1935) used cape syntektite for the same feature. kap bull 73ø-34 (73°44.1´n 23°50.5´w). cape on the north side of the mouth of moskusoksefjord. named by a.g. nathorst’s 1899 expedition for henrik johan bull [1844–1930]. nathorst’s ship, the antarctic, that he used for his 1898 expedition to spitsbergen and 1899 expedition to east greenland, had previously been used by the norwegian antarctic expedition led by bull in 1894–95. (cape bull). kap buxtorf 72ø-119 (72°53.0´n 25°43.4´w; map 4). cape west of the mouth of lumskebugten, se suess land. so named by eugène wegmann during the 1931–34 treårsekspeditionen after august buxtorf [1877–1969], a swiss structural geologist and stratigrapher who had mapped large areas of the swiss alps. (cape buxtorf.) kap börgen 75ø-22 (75°25.5´n 18°02.7´w; maps 2, 4). north cape of the island of shannon. named cap börgen by karl koldewey’s 1869–70 expedition after carl nicolay jensen börgen [1843– 1909], astronomer and geophysist at the observatory at leipzig. he was meteorologist and astronomer on koldewey’s 1869–70 expedition, and one of the original party to explore the island in august 1869. kap koner was also used for the cape in some of the expedition reports. (cape börgen, kap børgen.) kap canis major 80ø-107 (80°36.6´n 19°22.0´w; map 4). cape on the south side of inner ingolf fjord, opposite hjørnegletscher, kronprins christian land. named by john haller following explorations during lauge koch’s 1956–58 expeditions, after the constellation canis major. kap canis minor 80ø-106 (80°33.3´n 18°13.2´w). cape on the north side of central ingolf fjord, where the fjord narrows and is bordered by steep cliffs. named by john haller following explorations during lauge koch’s 1956–58 expeditions, after the constellation canis minor. kap carita 70ø (70°29.0´n 28°15.0´w). cape at the south entrance to vestfjord, the present renodde. the name was reported by helge g. backlund (in: koch 1955) as used by members of his group during the 1931–34 treårsekspeditionen. eklund (1944) also used this name in a report on mineralisation. kap carl ritter 76ø-2 (76°07.3´n 19°44.7´w; map 4). cape on the east coast of ad. s. jensen land. named cap carl ritter by karl koldewey’s 1869–70 expedition, for carl ritter [1779–1859], an influential german geographer. ritter was professor of geography at the university of berlin and first president of the berlin geo graphical society. the ‘carl-ritter-stiftung’ had made a grant towards publication of koldewey’s expedition narrative. a norwe gian hunting station south of the cape was sometimes known as carl ritterhytten – see olestua. (kap karl ritter.) kap christian 76ø-36 (76°36.5´n 18°35.7´w; map 4). south cape of lille koldewey, so named during the 1906–08 danmark-eks pedi tionen by christian bendix thostrup for his son christian (thostrup 2007). kap copeland 75ø-23 (75°20´n 18°55´w; maps 2, 4). nw cape of shannon. named cap copeland by karl koldewey’s 1869–70 expedition, after ralph copeland [1837–1905], a surveyor on kolde wey’s expedition and in part responsible for the astronomical section of the expedition narrative. in 1870 he was appointed astronomer to the observatory at parsonstown, ireland, and was from 1889 director of edinburgh observatory. kap copeland hytten 75ø (75°15.1´n 18°49.4´w). danish hunting hut on the west coast of shannon about 10 km south of kap copeland, built by nanok in september 1948. it has also been known as haraldsborg. (copelandshytten.) kap curly lillie 76ø-87 (76°39.5´n 18°30.6´w). south cape of the island maroussia, south of danmark havn. so named by the 1906–08 danmark-ekspeditionen. origin of name unknown. kap dalton [ittertivaa] 69ø-8 (69°24.7´n 24°04.0´w; map 3). bold headland on the northern part of the blosseville kyst. it was named cape dalton by william scoresby jr. during his 1822 voyage after john dalton [1766–1844]; a chemist and natural philosopher, dalton was the first to describe colour blindness. (kap daltonip nuua.) kap daly 75ø-14 (75°27.2´n 21°24.5´w; map 4; fig. 51). cape in the inner part of ardencaple fjord, between bredefjord and 216 smalle fjord. it was named cap daly by karl koldewey’s 1869–70 expedition, possibly after charles patrick daly [1816– 99], president of the american geographical society from 1864 to 1899 (j. løve, personal communication 2010). kap david gray 74ø-23 (74°58.0´n 18°26.6´w; maps 2, 4). south cape of the island shannon. named cap david gray by karl kolde wey’s 1869–70 expedition for captain david gray [1829–96], who had corresponded with august petermann on ice conditions off east greenland in 1872. david gray was one of the noted peter head whaling family. known as the ‘prince of whalers’, he made 43 voyages to the arctic from 1867 to 1890 in the eclipse, and had a total reported catch of 197 whales and 168,956 seals. (cape david gray.) kap david grayhytten 74ø-107a (74°59.0´n 18°23.7´w). danish hunting hut 2 km ne of kap david gray on the south coast of shannon. it was built by østgrønlandske fangstkompagni in 1923, and from 1929 used by nanok. it was also known as jægerly. in 1930 the hut was rebuilt by j. van hauen and a. hvidberg, but is now a ruin. a norwegian hut built at the same locality in 1952 was known as tåkeheim. (david gray hytten, kap david gray-hus, kap david grey hytten.) kap desbrowe 74ø-15 (74°38.3´n 18°19.8´w; map 4). se cape of lille pendulum. named cape desbrowe by douglas clavering in 1823 at the request of captain edward sabine, in honour of edward desbrowe, member of parliament for windsor and vice-chamberlain to queen charlotte. desbrowe had assisted sabine’s entry into the army. the present position corresponds to that of clavering’s description of a bold headland, although his map is inaccurate and the maps of koldewey’s 1869–70 expedition placed the name against the southernmost low cape. james wordie noted the discrepancy in 1929, and considered the name may have been intended for the sw cape below the mountain terrassebjerg. kap desbrowe hus 74ø (74°36.7´n 18°23.9´w). danish hunting hut on se sabine ø, about 4 km sw of kap desbrowe, built in the summer of 1921 by østgrønlandske fangstkompagni. it was later used by norwegian hunters who called it pendelbua. (kap des browe hytten). kap drygalski 79ø-27 (79°00.1´n 19°10.6´w; map 4). north cape of the small unnamed island north of achton friis ø at the front of zachariae isstrøm. the cape was named by the 1906–08 danmarkekspeditionen after erich von drygalski [1865–1949], a noted german geographer and geophysicist who led expeditions to greenland in 1891 and 1892–93, and the 1901–03 german south pole expedition. he was professor in geography at munich from 1906 to 1935. kap dufva 72ø-40 (72°40.2´n 24°42.6´w; map 4). cape in eastern lyell land. named by a.g. nathorst’s 1899 expedition after john hilmar dufva [b. 1864], a stockholm businessman who guaranteed a sum of 2500 swedish kronor for the expedition finances. a hunting hut thought to lie about 4 km west of kap dufva is now considered never to have been built (p.s. mikkelsen 1994, 2008). (cape dufva.) kap dundee [pukkitsivakajiip akinnarteqitaa] 69ø-60 (69°45.3´n 23°13.0´w). cape on the west side of the mouth of deichmann fjord, northern blosseville kyst. named by malcolm slesser during his 1969 expedition after the university of dundee, to which one of the expedition members (ian smart) was affiliated. dundee, the third city of scotland, dates from the late 12th century, while its university was founded in 1881. kap ehrenberg 74ø-69 (74°26.7´n 21°47.0´w). cape in eastern payer land where rudi bugt meets tyrolerfjord. it was named cap ehrenberg by karl koldewey’s 1869–70 expedition after christian gottfried ehrenberg [1795–1876], who contributed one of the chapters of koldewey’s narrative. he was one of the pioneers of microbiology and micropalaeontology. (cape ehrenberg, kapp ehren berg.) kap elisabeth 72ø-52 (72°54.3´n 24°48.5´w). ne cape of ella ø. named by a.g. nathorst in 1899, probably after his daughter elisabeth jane [b. 1885]. the norwegian hunting station 3 km south of the cape is known as maristua. (kap elizabeth.) kap ellen 77ø-32 (77°27.1´n 20°21.1´w). cape between helge g. backlund fjord and v. clausen fjord, west of skærfjorden, so named during the 1906–08 danmark-ekspeditionen. it may have been named by henning bistrup after his wife, ellen marie birgitte eigtved. kap eva 74ø-272 (74°09.8´n 22°12.8´w; map 4). cape on the north side of jordanhill opposite kap adam. named as kapp eva on the 1932a nsiu map. the two capes were evidently named after adam and eve. kap ewart 69ø-9 (69°21.0´n 24°26.0´n; map 3). cape on the north blosseville kyst. named by william scoresby jr. in 1822 as cape ewart, probably after peter ewart [1767–1842], an engineer, and owner of a cotton mill in manchester. kap fennia 71ø-114 (71°16.7´n 21°52.5´w). cape in east liverpool land west of trekanten. so named by helge g. backlund during the 1931–34 treårsekspeditionen, after his homeland finland. kap fletcher 71ø-13 (71°37.1´n 22°06.0´w; map 4). cape on the east coast of canning land, named cape fletcher by william scoresby jr. in 1822 after an edinburgh friend, possibly john fletcher [1792–1836]. scoresby’s original cape was probably 3–4 fig. 51. kap daly separates smallefjord at left from bredefjord to the right. the mountain between the two fjords is known as storborgen. ejnar mik kelsen gletscher is just visible at the inner end of bredefjord. storborgen bredefjord sm alle fjo rd kap daly 217 km further south, the position used on koch’s (1902) map. the present position was chosen as being ‘more natural’ by the place name committee. (kap fletscher.) kap franklin 73ø-7 (73°15.0´n 22°12.7´w; maps 3, 4). cape on the east point of gauss halvø. it was named by william scoresby jr. in 1822 as cape franklin after john franklin [1786–1847], a noted arctic explorer, whose last expedition to discover the nw passage was lost with all hands. this calamity led to an important series of search expeditions in the canadian arctic, which franklin had earlier explored during two overland expeditions. karl koldewey’s 1869–70 expedition maps and 1888 danish charts place the name too far north, while payer (1876) used the correct (present) position. a norwegian hut 7 km north of the cape was sometimes known as kap franklin – see franklin stranda. (kapp franklin.) kap freuchen 76ø-132 (76°21.0´n 23°41.4´w; map 4). cape-like peninsula in dronning louise land between budolfi isstrøm and l. bistrup bræ, named by j.p. koch’s 1912–13 expedition after peter freuchen, their companion during the 1906–08 danmarkekspeditionen. peter freuchen [1886–1957] was stoker and assistant meteorologist on the 1906–08 danmark-ekspeditionen. in 1910 he accompanied knud rasmussen to thule and helped establish the trading station, of which he was manager until 1919. he took part in the 1st and 5th thule expeditions, and travelled extensively in the arctic. he is best known as the author of stories of eskimo (inuit) life based on his experiences. kap givagt 74ø (74°48.7´n 20°39.8´w). cape on the west side of kuhn ø. the name is seen in reports by helge g. backlund on his work during the 1931–34 treårsekspeditionen (in: koch 1955), and warns of the numerous, dangerous sandbanks around the cape (giv agt = beware). kap gladstone 71ø-11 (71°31.4´n 21°53.2´w; map 4). bold headland forming the northern termination of liverpool land. named cape gladstone by william scoresby jr. in 1822 (fig. 3) as a compliment to john gladstone [1764–1851], an enterprising liverpool merchant and member of parliament. kap godfred hansen 71ø-198 (71°26.8´n 21°42.1´w; map 4; see also fig. 72). peninsula on the east coast of north liverpool land. the name was originally used by henning bistrup on his coast profiles drawn in 1923 during the drift of the teddy, but was not approved until 1939. see also godfred hansen ø. kap graah – aa is treated as å in danish – see after kap greville. kap graham 69ø-2 (69°59.0´n 22°29´w; map 4). cape sw of kap brewster between kap russel and kap pillans. it was named by william scoresby jr. in 1822, although the name does not occur on his chart. from his table of latitudes and longitudes it is clear that scoresby’s cape graham was intended for a cape west of steward ø, corresponding probably to the present akinnarteqitaa. its current position derives from its order of listing in scoresby’s text (scoresby 1823, p. 231). all three capes were named after professors at the university of edinburgh. possibly named after robert graham [1786–1845], first regius professor of botany 1818–1820, and founder of the edinburgh botanical gardens. kap greg 70ø-248 (70°56.9´n 21°37.7´w; maps 3, 4). headland, almost an island, on the east coast of liverpool land. named cape greg by william scoresby jr. in 1822 out of respect and regard to samuel greg [1758–1834] of quarry bank. a hunting hut was built on the low col west of the cape by scoresbysund municipality. kap greville 71ø-29 (71°29.9´n 22°05.8´w; map 4). cape in north liverpool land. william scoresby jr. named cape greville in 1822, with several other promontories after different friends chiefly resident in edinburgh. robert kaye greville [1794–1866], was a noted botanist, and like scoresby a member of the wernerian society and the royal society of edinburgh. (cape grewille). kap graah 73ø-22 (73°14.3´n 23°12.6´w; maps 3, 4). east cape of gunnar andersson land, the northern part of ymer ø. named cap graah by karl koldewey’s 1869–70 expedition for wilhelm au gust graah [1793–1863], a danish naval officer, especially noted for his expedition to se greenland in 1828–31 which mapped the east coast of greenland up to 65°18´n. (cape graah, kap gråh.) kap h.n. andersen 80ø-1a (80°02.1´n 17°16.2´w; maps 1, 4). ne cape of hovgaard ø. so named by the 1906–08 danmark-eks peditionen after hans niels andersen [1852–1937], one of the founders and a director of the østasiatiske kompagni, then den mark’s largest shipping company. the company had contributed 10 000 danish kroner to the expedition finances. (cape h.n. andersen.) kap hamburg 74ø-34 (74°42.4´n 20°03.8´w; map 4). south cape of kuhn ø. named cap hamburg by karl koldewey’s 1869–70 expedition, probably for the german city of hamburg, the base for one of the expedition’s supporting committees, the ‘hamburger comité fur die nordfahrt v. 1869’. the senate of hamburg made a large contribution to the expedition finances. a norwegian hut 3 km west of the cape, sometimes known as kap hamburg hytten, is more commonly called furnes. (kap hamborg.) kap harry 72ø-53 (72°46.5´n 24°52.3´w; map 4). se cape of ella ø. named by a.g. nathorst’s 1899 expedition, probably after his son harry johan hjalmar nathorst [1882–1938], who became a mining engineer. many features in the vicinity were named by nathorst after members of his family. (cape harry.) kap hartlaub 74ø-39 (74°42.5´n 18°18.6´w). cape in ne lille pendulum. named cap hartlaub by karl koldewey’s 1869–70 expedition after carl johan gustav hartlaub [1814–1900], an ornithologist noted for his studies of african birds. hartlaub together with moritz lindeman edited koldewey’s narrative and scientific reports for publication. kap hedlund 72ø-407 (72°43.6´n 26°11.2´w; map 4; fig. 52). prominent cape in nw lyell land at the mouth of rhedin fjord. named by a.g. nathorst’s 1899 expedition, possibly after both sven adolf hedlund [1821–1900] and henrik hedlund [1851– 1932]. both were publishers, and both had made financial contributions to nathorst’s expedition. (cape hedlund.) kap hedlund hytten 72ø (72°43.1´n 26°10.5´w). hut built by sirius in 1964 in the bay on the east side of kap hedlund. it replaced the norwegian hut on the same site known as kapp hedlund hytta or rimhytten. kap helgoland 76ø-8 (76°43.5´n 19°05.6´w; map 4). northern cape of store koldewey. named cap helgoland by karl koldewey’s 1869–70 expedition after the island off the nw coast of germany. kap helgoland hytten 76ø (76°43.2´n 19°06.3´w). hut built by danmarkshavn weather station personnel in 1965 in a small bay near kap helgoland. it is also known as aldersro. kap hendil 73ø-591 (73°28.5´n 27°20.5´w; map 4). se cape of louise boyd land. this is one of the names found on the 1932 edition of the geodætisk institut 1:1 million scale map, drawn from lauge koch’s aerial observations during the 1931–34 treårs ekspeditionen. it was named after leif hendil [1898–1961], a journalist with the danish newspaper ekstrabladet, who had helped attract private financial support for lauge koch’s expeditions. kap herschell – see herschellhus. kap hewitt 71ø-8 (71°24.5´n 21°40.7´w; maps 3, 4). cape on the east coast of liverpool land, named by william scoresby jr. in 1822. it was one of several names in the north part of liverpool land which scoresby gave for different unspecified friends chiefly resident in manchester. scoresby probably intended his cape hewitt to correspond to that now known as kap godfred hansen, a position also used on unpublished profiles drawn by henning bistrup. however, scoresby’s chart and early danish maps are poor in northern liverpool land, and misplacement of named features is not surprising. kap hilding 71ø-118 (71°16.6´n 21°46.3´w). cape on the west side of trekanten, east liverpool land. so named by helge g. back lund during the 1931–34 treårsekspeditionen after his youngest 218 son. officially it is stated to be in the sense of ‘hilding’ (= giant or warrior). kap hodgson [kiammut nuukajia] 70ø-206 (70°33.5´n 21°30.3´w; maps 3, 4). prominent cape in se liverpool land. named cape hodgson by william scoresby jr. in 1822 after his friend adam hodgson. kap hold with hope station – see kap broer ruys station. kap hooker 70ø-95 (70°27.0´n 23°16.3´w; map 4). the south point of jameson land was named cape hooker by william scoresby jr. in 1822 after william jackson hooker [1765–1865], professor of botany at the university of glasgow from 1821, and from 1841 director of kew gardens. hooker contributed the list of plants that appeared as appendix ii in scoresby’s narrative. ryder (1895) observed that the term cape does not fit very well because of the smooth rounding of the low and flat coastline, and the name has been placed as far north as the vandreblokken. in 1965 a location at the mouth of fynselv was selected by the place name committee. (cap hooker.) kap hope [ittaajimmiit] 70ø-287 (70°27.5´n 22°20.9´w). green landic village east of kap hope (see below) in south liverpool land. the 1924–25 expedition that founded scoresbysund built two houses here in 1924–25, and it has been continuously occupied until the late 1980s. the population in 1987 was reported as 41. official ministry reports used the greenlandic names itterajivit and illukasiit for the settlement up to 1987, although the place name committee had substituted igtâjingmit (ittaajimmiit) for igterajivit (itterajivit) in 1978 to comply with the usage of the inhabitants. kap hope [noorajik kangitteq] 70ø-286 (70°27.7´n 22°22.9´w; maps 3, 4). sw point of liverpool land, named cape hope by william scoresby jr. in 1822 out of respect to samuel hope of everton. the settlement east of the cape is known as kap hope [ittaajimmiit] (see also above). kap hovgaard 72ø-71 (72°41.2´n 22°37.6´w; fig. 14). cape on southern geographical society ø, west of nordenskiöld ø. the name was given by j.m. wordie’s 1926 expedition, originally as cape hovgaard, to commemorate the danish member of the vega expedition through the ne passage. andreas peter hov gaard [1853–1910], a lieutenant in the danish navy, was in charge of the magnetic and meteorological work of the vega expedition. see also hovgaard ø. (kap hovgård, kapp hovgaard). kap humboldt 73ø-5 (73°05.7´n 23°01.2´w). eastern cape of ymer ø. william scoresby jr. named cape humboldt in 1822 in compliment to the celebrated traveller friedrich heinrich alex ander, freiherr von humboldt [1769–1859]. humboldt was noted for his travels in central and south america betwen 1799 and 1804. the ‘cape’ sighted by scoresby was probably either the present bontekoe ø (which he placed farther north), or possibly celsius bjerg. koldewey’s maps (verein für die deutsche nordpolarfahrt in bremen 1873–74) moved the name to a cape west of bontekoe ø on the south side of kejser franz joseph fjord, a position more precisely defined by nathorst (1900) as the east cape of ymer ø. kap humboldt fangststation – see humboldt. kap hynæs – see kapp agnes. kap høegh [ukaleqarteq] 70ø-226 (70°43.4´n 21°33.3´w; map 4). east cape of sandbach halvø, south liverpool land. the name first appeared on a map compiled by janus sørensen (sørensen 1928), and was evidently given for henrik høegh, manager of the scoresbysund colony from 1926. the spelling of the original map, kap höegh, has survived on many published maps. a hunting hut was built on the low col west of the cape for the use of hunting parties from scoresbysund. kap ingrid 77ø-73 (77°38.0´n 20°21.2´w; map 4). cape east of the mouth of campanulavigen, inner skærfjorden. named by david malmquist during the 1931–34 treårsekspeditionen after his sister, ingrid madsen. kap isabelle 77ø-5 (77°44.5´n 19°08.3´w; map 4). se cape of gamma ø, on the north side of the mouth of orléans sund. named by the duke of orléans in 1905 as cap isabelle, probably after his mother, isabelle de montpensier [1848–1919], countess of paris. kap james 73ø-15 (73°53.1´n 20°18.3´w; map 4). ne cape of home forland on the south side of gael hamke bugt. it was named cape james by douglas clavering in 1823 after his friend james smith [1782–1867], who wrote the introduction to clav ering’s (1830) narrative of his 1823 expedition. smith was a geologist and writer, and generally known as ‘smith of jordanhill’. see also kap mary. (kapp james.) kap james hytten 73ø (75°53.1´n 20°18.3´w). norwegian hunting hut built by møre grønlandsekspedition in 1930 on the nw side of kap james. it was originally known as röbeckstua. kap jarner 76ø-119 (76°38.0´n 22°08.2´w; map 4). cape on the south side of borgfjord, named by j.p. koch’s 1912–13 expedition after the geologist of the 1906–08 danmark-ekspeditionen, hakon høeg jarner [1882–1964]. jarner was trained as an architect, and for most of his career worked as a factory inspector. see also jarners kulmine. (jarnerhöfdi.) kap knut söderström röhss fjord rhedin fjord kempe fjord dickson fjord kap bayard suess land kap beijer kap hedlund fig. 52. view northwards over kap hedlund, where kempe fjord divides into three fjords: rhedin fjord, röhss fjord and dickson fjord, with the prominent capes kap knut söderström, kap bayard and kap beijer. suess land is in the background. the john haller photograph collection, geus archive. 219 kap jona 74ø (74°36.7´n 18°23.9´w). danish hunting hut built in 1921 for østgrønlandsk fangstkompagni. jonas karlsbak came across the hut in 1928 and called it kap jona after one of his daughters, and it has subsequently figured in norwegian hut lists under this name or as pendelbua. kap jones 71ø-1 (71°07.3´n 21°43.3´w; maps 3, 4). cape on the north side of the entrance to storefjord, liverpool land. it was named cape jones by william scoresby jr. in 1822 in compliment to john jones [1791–1889], a successful evangelical minister, who in 1815 became the first incumbent of st. andrews church, liver pool. h.g. backlund used kap anatektite for the same feature during his 1933 explorations of liverpool land. kap jungersen 80ø-11 (80°36.5´n 16°05.2´w; maps 1, 4). south cape of amdrup land. so named by the 1906–08 danmarkekspeditionen for hector frederik estrup jungersen [1854–1917], a danish zoologist who took part in the deep-sea hydrographical 1895 and 1896 voyages around iceland with the ingolf, and was professor of zoology at the zoological museum in copenhagen from 1899. he was a member of the committee of the 1906–08 danmark-ekspeditionen. (kap-jungersen-fjæld.) kap jørn 71ø-62 (71°37.0´n 27°26.0´w). cape in ne hinks land on the north side of the mouth of flyverfjord. it was mapped by lauge koch from the air in 1932, during the 1931–34 treårs ekspeditionen, and named after the son of victor petersen, pilot of koch’s seaplane. kap klinkerfues 75ø-12 (75°17.4´n 20°38.0´w). cape on the north side of the mouth of ardencaple fjord. named by karl koldewey’s 1869–70 expedition after ernst friedrich wilhelm klinkerfues [1827–1884], a german astronomer who was director of göt tingen observatory. he was a good friend of koldewey’s. (cap klinkerfues, cape klinkerfues, kap klingafus.) kap knut söderström 72ø-408 (72°44.0´n 26°18.9´w; fig. 52). cape in eastern gletscherland between röhss fjord and rhedin fjord. named by a.g. nathorst’s 1899 expedition for knut söder ström, a supporter of the expedition who had donated generous quantities of wines and cognac. (cape knut soderstrom.) kap koefoed 78ø-7 (78°29.5´n 18°23.6´w; maps 1, 4). east cape of the southernmost island of the franske øer. named by the duke of orléans in 1905 as cap koefoed, after einar laurentius koefoed [1875–1963], a zoologist who was the expedition biologist. the cape has been placed farther north on some maps. kap kolthoff 73ø-35 (73°43.3´n 24°02.0´w; map 4). nw cape of gauss halvø at the entrance to moskusoksefjord. named by a.g. nathorst’s 1899 expedition after gustaf isak kolthoff [1845– 1913], a swedish zoologist, and conservator at the university of uppsala from 1878 to 1912. nathorst described him as a valued friend and companion on two polar voyages (spitsbergen in 1898 and east greenland in 1899). in 1900 kolthoff led his own zoological expedition to east greenland. (cape kolthoff.) kap koner 75ø (75°25.5´n 18°02.7´w). name used for the ne point of shannon in the geology section of karl koldewey’s narrative (verein für die deutsche nordpolarfahrt in bremen 1873–74) of his 1869–70 expedition. it was named after wilhelm david koner [1817–1887], a german geographer and librarian at the university of berlin. elsewhere in koldewey’s maps and text it is replaced by cap börgen (now kap börgen). kap kraus 73ø-19 (73°47.4´n 20°18.3´w; map 4). se cape of home forland, named as cap kraus by karl koldewey’s 1869–70 expedition. it was probably named after gregor konrad michael kraus [1841–1915], professor and director of the botanical gardens at erlangen. he contributed a chapter on driftwood to koldewey’s scientific reports (j. løve, personal communication 2010). (cape kraus.) kap kuhre 76ø (76°34.1´n 19°03.6´w). cape on the south side of the mouth of berg fjord, store koldewey. the name is used in den grønlandske lods (1968), and is an adaption of kap kuré, a name proposed by j.g. jennov, and given for the captain of the gefion during the 1932 expedition. neither version of the name is approved. kap kuré – see kap kuhre. kap lagerberg 72ø-37 (72°31.4´n 24°39.5´w; map 4). cape in east lyell land. named by a.g. nathorst’s 1899 expedition, possibly after carl sven axel lagerberg [1822–1905], a count and army general, reported as a popular swedish figure. (cape lagerberg.) kap lagerberg hytten – see beinhaugen. kap lapparant 73ø-624 (73°14.4´n 26°10.5´w). south cape of andrée land, so named during the 1931–34 treårsekspeditionen by eugène wegmann in the forms cape lapparant and cape lapparent. it is said to have been given for several french mineralogists and geologists: albert auguste de lapparent [1839–1908], albert felix de lapparent [1905–1975] and jacques de lapparent [d. 1949]. kap laplace – see laplace huset. kap leslie [ilimananngip nunaa] 70ø-48 (70°39.2´n 25°16.4´w; maps 3, 4). eastern cape of milne land, which william scoresby in 1822 named as cape leslie in compliment to john leslie [1766– 1832]. scoresby attended professor leslie’s mathematics lessons at the university of edinburgh in 1808. kap li 77ø-14 (77°21.0´n 19°48.1´w; map 4). cape at the south side of the the mouth of c.f. mourier fjord in sw skærfjorden. so named by david malmquist during the 1931–34 treårs ekspedi tionen for li hadders [b. 1913], whom he married in 1935. kap li hytten – see knuthsminde. kap lister 70ø-340 (70°29.5´n 21°32.8´w; map 4). cape in se liverpool land. william scoresby jr. named cape lister in 1822 after a friend, the revd lister. the cape was one of scoresby’s landing places during his 1822 voyage. kap louise 77ø-6 (77°42.5´n 19°11.1´w; map 4). cape in ne stormlandet on the south side of the mouth of orléans sund. named cap louise by the duke of orléans in 1905, possibly after louise [d. 1850], a sister of his grandfather ferdinand who was married to léopold 1 of belgium. on one of the expedition maps the same cape is named cap de guise. kap mcclintock 72ø-72 (72°40.7´n 21°56.1´w; maps 3, 4). se cape of geographical society ø. the name was given by j.m. wordie’s 1926 expedition to the point opposite kap parry to commemorate the arctic explorer, leopold m’clintock. named originally in the form cape mcclintock or c. mc.clintock, it was adopted on nsiu maps in the form kapp mac clintock and on danish maps as kap mac clintock, the usual danish convention for scottish names of this type. sir francis leopold m’clintock [1819–1907], a british naval officer and explorer, was most noted for his 1857–59 voyage in the fox, which found the cairn record revealing the fate of the 1845 franklin expedition. kap macclintock hytten 72ø (72°40.9´n 22°02.1´w). sirius hut erected in 1956 on a small peninsula about 3 km west of kap mc clin tock. it is also known as valmuehytten. kap mackenzie 72ø-17 (72°53.8´n 21°53.8´w; maps 3, 4). ne cape of geographical society ø. the name mackenzie island first appeared on the 1872 edition of british admiralty chart 2282 together with franklin island. white (1927) suggested the two names owe their origin to a mistake by the draughtsman, who may have had mackenzie bugt and kap franklin in mind when engraving the copper plate. wordie found the supposed island to be a cape in 1926, and named it cape mackenzie. kap madelaine 73ø-697 (73°19.7´n 26°44.0´w). prominent cape in sw andrée land, on the ne side of isfjord. named by john haller following explorations during lauge koch’s 1949–51 expeditions. kap margrethe 72ø-270 (72°53.4´n 24°47.8´w). minor cape on ne ella ø. named by john w. cowie during work carried out from 1949 to 1954 on lauge koch’s geological expeditions. it is said to have been given for the eldest of the danish princesses, margrethe 220 alexandrine þórhildur ingrid [b. 1940], the eldest daughter of frederik ix of denmark, who became queen margrethe ii of denmark in 1972. kap marie dijmphna 80ø-30 (80°04.6´n 18°02.7´w). cape on northern hovgaard ø, on the south side of dijmphna sund west of kap povl. so named by the 1938–39 mørkefjord expedition after eigil knuth’s great grandmother marie dijmpha [1813–1876]. the ship dijmphna, used for the danish expedition to the kara sea in 1882–83, was christened by knuth’s mother, marie gàmel. kap marie valdemar 77ø-12 (77°15.8´n 18°20.9´w; maps 2, 4). cape in northern germania land, named in 1905 as cap marie waldemar by the duke of orléans. the original cape was the present kajkap farther west, and the name was accidently transferred to the present location by the 1906–08 danmark-ekspe ditionen, who used it extensively in their reports before the error of position was discovered. (cape marie-valdemar.) kap martha hytten – see kapp martha. kap mary 74ø-22 (74°09.7´n 20°11.7´w; map 4). cape on eastern clavering ø, on the north side of gael hamke bugt. named cape mary by douglas clavering in 1823 after the wife of his friend james smith. see also kap james. mary wilson [d. 1847] had married smith in 1809. two hunting huts built at the cape were known as maryhuset and christianshavn. (kapp mary.) kap maurer 74ø-26 (74°51.5´n 19°44.4´w; map 4). cape on east kuhn ø. it was named cap maurer by karl koldewey’s 1869–70 expedition after professor konrad von maurer [1823–1902]. he studied natural sciences and law, and contributed a chapter on the exploration of greenland to koldewey’s expedition narrative (j. løve, personal communication 2010). kap maurer hytten 74ø (74°48´n 19°51´w). danish hunting hut about 8 km south of kap maurer on the east coast of kuhn ø, built by nanok in 1930. it was also known as jennovshåb. (kap maurer hytten, maurer-hytten.) kap mcclintock, kap mcclintock hytten – ‘mc’ is treated as ‘mac’ – see above before kap mackenzie kap menelik 77ø-146 (77°05.3´n 20°57.3´w; map 4). cape on the south side of sælsøen. named during lauge koch’s 1956–58 expeditions by john haller after menelik ii, also known as sahle miriam [1844–1913], one of ethiopia’s greatest rulers. he is said to have played a role for some members of the 1906–08 danmarkekspeditionen. kap mérite 78ø-3 (78°14.5´n 18°50.0´w; map 4). east cape of the island stigbøjlen. named by the duke of orléans in 1905 as cap mérite, after édouard mérite [1867–1941], painter and naturalist on the expedition. kap mohn 73ø-507 (73°11.6´n 25°45.2´w; map 4). western cape of ymer ø. named during karl koldewey’s 1869–70 expedition, although the name is only found in the narrative of payer (1876) in the form insel mohn. henrik mohn [1835–1916] was a norwegian meteorologist, founder and director of the meteorological institute in christiania (now oslo), and had corresponded with the expedition committee and payer. mohn encouraged norwegian sealer captains to make geographical and meteorological observations during their voyages, and their results were published in peter manns mitteilungen. a.g. nathorst observed in 1899 that the island depicted by payer was joined by a low promontory to another island (insel petersen), and moved both names to western capes of ymer ø. (mohn insel, cape mohn.) kap montpensier 77ø-2 (77°51.5´n 17°36.6´w; maps 1, 2, 4). northern cape of île de france (from 2004 qeqertaq prins henrik). named by the duke of orléans in 1905 as cap montpensier, probably after his mother, isabella de montpensier [d. 1919]. kap moorsom 72ø-4 (72°10.5´n 22°06.5´w; map 4). short pro montory on se traill ø, named cape moorsom by william scores by jr. in 1822 out of respect to richard moorsom jr. of whitby. kap mosle 75ø-21 (75°02.4´n 20°23.0´w; map 4). nw cape of kuhn ø. named by karl koldewey’s 1869–70 expedition as cap mosle, after alexander georg mosle [1827–1882]. he was president of the ‘bremisches comité für die zweite deutsche nord polar fahrt’, one of the expedition’s principal supporting orga ni sations. kap mæchel 72ø-29 (72°23.5´n 25°15.5´w; maps 4, 5; see also fig. 61). cape between forsblad fjord and alpefjord, named by a.g. nathorst’s 1899 expedition. the name appears on charts in both swedish (1900) and english (1901) editions of nathorst’s narrative in the form ‘mæchel’, but appears in the index of the swedish edition as kap maechel. it was evidently named after captain e. maechel of the swedish royal navy who had assisted nathorst in his choice of ships for his voyages to spitsbergen and east greenland. (cape maechel, cape moechel.) kap mæchelhytten – see mæchel-stua. kap möbius – see möbius bjerg. kap nansen 79ø-13 (79°10.7´n 17°46.3´w; map 1, 4). north cape of the largest of the norske øer. named by the 1938–39 mørke fjord expedition after the norwegian arctic explorer fridtjof nansen [1861–1930], who was noted especially for his crossing of the inland ice of greenland in 1888, and his drift across the arctic ocean with the fram in 1893–96. kap nax 77ø-30 (77°32.8´n 19°56.5´w; map 4). se cape of c. silfverberg ø, so named by the 1906–08 danmark-ekspeditionen. origin of name unknown. kap negri 75ø-20 (75°03.0´n 20°37.9´w; map 4). cape on the south side of the mouth of grandjean fjord. named cap negri by karl koldewey’s 1869–70 expedition, after baron christoforo negri [1809–96]. an italian geographer, he was founder and first president of the ‘reale società geografica italiana’, and a supporter of the expedition. a norwegian hunting hut built at the cape by sigurd tolløfsen’s expedition, and sometimes known as kap negri hytten, is now in poor condition; it is better known under the names city hytta and vedethytten. kap neumayer 74ø-43 (74°40.7´n 18°51.9´w; map 4). northern cape of sabine ø. named cap neumayer by karl koldewey’s 1869–70 expedition after georg balthasar von neumayer [1826– 1909]. a german meteorologist and oceanographer, he was founder of ‘deutsche seewarte hamburg’ (german naval obser va tory, hamburg), and a promoter of polar research. the success of the first international polar year 1882–83 is attributed in large part to neumayer. (cape neumayer, kap neumayr.) kap niels 76ø-19 (76°23.3´n 21°35.2´w; map 4). cape on the east coast of rechnitzer land, so named by henning bistrup during the 1906–08 danmark-ekspeditionen. possibly named after niels baron juel-brockdorff, a colleague at the marine cadet school in 1898 (j. løve, personal communication 2009). kap niels hytten 76ø (76°25.5´n 21°37.7´w). norwegian hunting hut at the peninsula north of kap niels, ne rechnitser land, built in august 1933 by john giæver’s expedition. hunters had assumed this more prominent cape was kap niels, and as it has no other name the name is still often used. kap nielsen fjeld – see kai nielsen fjeld. kap norge – see kapp norge. kap oetker 74ø-77 (74°15.3´n 21°59.8´w). cape on sw clavering ø. named cap oetker by karl koldewey’s 1869–70 expedition after friedrich oetker [1809–81], a german author and lawyer. huts at and se of the cape have been known as kap øtker hytten (see nes-odden) and kapp oetker. (cape oetker, kapp oetker.) kap olga – see cap holcha. kap oswald 72ø-51 (72°53.0´n 25°08.1´w). cape on nw ella ø. so named by a.g. nathorst’s 1899 expedition, possibly for oswald heer [1809–1883] (see also kap oswald heer), or more probably for a member of nathorst’s own family. (cape oswald). kap oswald heer 75ø-8 (75°32.8´n 19°26.3´w; map 4). relatively elevated section of the east coast of hochstetter forland, with the 221 appearance of a cape in the field, although it is not particulary prominent on a map. named cap oswald heer by karl koldewey’s 1869–70 expedition after oswald heer [1809–83], a noted swiss botanist and geologist, professor in zurich from 1852 to 1882. he was an expert on arctic fossil floras, and contributed a section on fossil plants to koldewey’s narrative. (cape oswald heer, c. heer, c. osvald heer.) kap oswald heerhytten 75ø-98 (75°30.5´n 19°22.8´w). danish hunting hut about 4 km south of kap oswald heer, built by nanok in may 1931, and rebuilt in 1932 and 1933. (kap oswald heer hytten, oswald heer hytten.) kap ovibos 73ø-515 (73°33.1´n 24°24.1´w; map 4). se cape of strindberg land. so named by a.g. nathorst’s 1899 expedition for the musk ox (ovibos muschatus), of which he saw five at the cape. (cape ovibos.) kap oviboshytten 73ø (73°32.9´n 24°25.0´w). norwegian hut on the south side of kap ovibos, built by arktisk næringsdrift in sep tember 1933. it was originally known as solheim. (ovibos.) kap palander 72ø-61 (72°37.4´n 22°29.8´w). cape on eastern traill ø on the south side of vega sund. so named by a.g. nat horst’s 1899 expedition after adolf arnold louis palander af vega [1842–1920], baron, swedish admiral, explorer and politician. palander made several polar voyages, most notably through the ne passage and around asia as commander of the vega with n.a.e. nordenskiöld. (cape palander.) kap pansch 75ø-27 (75°09.4´n 17°24.4´w; maps 2, 4). cape on eastern shannon. named cap pansch by karl koldewey’s 1869–70 expedition after adolph georg pansch [1841–1887], the expedition doctor. he was professor of botany at zurich from 1852 to 1882, and contributed many of the narrative sections to koldewey’s book of the expedition (j. løve, personal communication 2010). kap parry 72ø-9 (72°24.0´n 21°56.8´w; maps 3, 4). cape on eastern traill ø, 609 m high. william scoresby jr. gave the name cape parry in 1822 to a bold headland on the north side of mountnorris fjord, in honour of captain william edward parry [1790–1855]. parry was noted for three voyages in search of the nw passage, in 1819–20, 1821–23 and 1824–25, and for an attempt to reach the north pole by boat in 1827. kap payer 73ø-567 (73°11.0´n 26°27.8´w; map 4). cape on the south side of kejser franz joseph fjord, north of payer tinde. the name was used first by lauge koch’s 1926–27 expeditions in the form cape payer, although judging from the description in koch (1930) for a less conspicuous cape 18 km east of the present location. see also payer tinde. kap peschel 76ø-3 (76°14.8´n 19°59.0´w; map 4). ne cape of ad. s. jensen land, south of roon bugt. named cap peschel by karl koldewey’s 1869–70 expedition after oskar peschel [1826–75], a german geographer who was professor at leipzig. one of the horse-sledges used by j.p. koch’s 1912–13 expedition is deposited on a small island off the cape (see slædeøen). a norwegian hunting hut west of the cape is sometimes known as kap peschelhytten (see strømsbukta). (peschelkap, kap peschell.) kap petersen 73ø-508 (73°23.9´n 25°17.5´w). western cape of gunnar andersson land, nw ymer ø. it was named petersen insel during karl koldewey’s 1869–70 expedition, although the name only occurs in the narrative of payer (1876). the origin of the name is uncertain, but it is likely to have been given for a norwegian scientist as are three other names only found on payer’s maps (the present kap mohn, broch øer and kjerulf fjord). as the supposed island did not exist, a.g. nathorst transferred the name to the cape in 1899. see also kap mohn. kap peterséns 72ø (72°25.3´n 24°37.0´w; map 5). cape on the sw side of kong oscar fjord at the mouth of segelsällskapet fjord. it was named by a.g. nathorst’s 1899 expedition, probably after carl justus frederik af petersens [1851–1925], a contemporary of nathorst’s at the university of lund who became notable as head of the university library. in his published maps nathorst distinguishes between kap peterséns and kap petersén (the latter a cape on ymer ø, spelt without the final ‘s’). this practice was followed on the maps of many subsequent explorers (e.g. wordie 1930a, b). kap peterséns figured on official danish maps for many years (the accent was added in 1935 by the place name committee as an aid to pronunciation), and is a reference locality often used in geological, botanical and climbing publications. it is also the type locality for a formation of the eleonore bay supergroup. the name was dropped from official danish maps in 1963, following allegations of confusion with kap petersen by knud lauritzen, the shipping magnate. however, the name continues to be used, both for the cape and the norwegian hunting station se of the cape (see below). (cape petersens, kap petersen, kapp pettersens.) kap peterséns 72ø (72°25.0´n 24°33.8´w). norwegian hunting station 2 km se of kap peterséns, built in 1930 by the møre expedition, and manned in the periods 1930–39 and 1951–59. the original name was sunnmørsheimen, but it is still generally known as kap peterséns, despite attempts to supress the name (see above). the station was regularly maintained and used by sirius after 1960, and was restored by nanok in 1997 and 1998. (kapp petersens.) kap philip broke 74ø-13 (74°55.8´n 17°36.9´w; map 4). south ernmost cape of shannon. named by douglas clavering in 1823 as cape philip broke for the commander of the frigate shannon under whom he had served as midshipman. sir philip bowes vere broke [1776–1841] had been appointed captain of the shannon in 1806, and was most noted for his capture of the chesapeake in 1813. a depot hut was built adjacent to the cape in 1901 (see below) for the baldwin-ziegler expedition. (kap philipp broke). kap philip broke 74ø (74°56.1´n 17°39.3´w). distinctive eight-sided hut just west of kap philip broke, and known by the same name. originally built as a depot hut for the 1901 baldwin-ziegler expedition, it was used as a refuge hut by members of the 1906–08 danmark-ekspeditionen and the 1909–12 alabama expedition, and later as a hunting hut by østgrønlandske fangstkompagni from 1920 to 1924 and nanok from 1929 to 1930. in 1930 the hut was transferred to norwegian ownership, but reverted to danish ownership in 1969 when all the other norwegian huts and stations in east greenland were taken over by the danish state. kap philippe 77ø-3 (77°36.5´n 17°45.9´w; maps 1, 2, 4). se cape of île de france (from 2004 qeqertaq prins henrik), named in 1905 as cap philippe. the name was given for philippe duke of orléans at the suggestion of his companions on the 1905 expedition. see also hertugen af orléans land. a cairn was built here on 29 july 1905, the record being recovered in 1988 by eigil knuth. (kap phillipe, isle de philippe.) kap pillans [immikkeerterajivit iliverta] 69ø-3 (69°56.7´n 22°35.3´w; map 4). cape sw of kap brewster. named cape pil lans by william scoresby jr. in 1822 after james pillans [1778– 1864], a scottish educational reformer who was professor of hu man ity and laws at the university of edinburgh from 1820 to 1860. kap povl 80ø-1 (80°04.6´n 17°34.6´w; map 4). ne cape of hovgaard ø, so named during the 1906–08 danmark-ekspedi tionen by j.p. koch after povl hammershøj [1905–61], the infant son of a friend. povl hammershøj became a major-general and military attache. (kap paul). kap quist hytten 76ø (76°43.3´n 18°32.2´w). hut on the east side of kap bismarck, se germania land. it was built in 1951 for dan marks havn weather station by steen malmquist. it is now a ruin. kap récamier 77ø-11 (77°23.2´n 19°56.7´w; map 4). cape on the north side of the mouth of c.f. mourier fjord. it was named by the duke of orléans in 1905 as cap récamier after joseph récamier [1774–1852], surgeon on the expedition and chief physician at the hôtel-dieu de paris, the oldest hospital in paris. kap reinhardt 75ø-13 (75°16.7´n 20°54.9´w). cape on the sw side of ardencaple fjord, north of the mouth of kildedal. named 222 by karl koldewey’s 1869–70 expedition as cap reinhard or cap reinhardt, probably after johannes theodor reinhardt [1816– 1882], the director of the natural history museum in copenhagen and professor at the university from 1865; he had been consulted on zoological questions by the expedition committee. kap rink 75ø-9 (75°07.9´n 19°36.7´w; maps 2, 4). south cape of hochstetter forland. named cap rink by karl koldewey’s 1869– 70 expedition after heinrich johannes rink [1819–93], a danish geologist, greenland explorer and administrator. he had corresponded with the expedition committee. the danish hunting station built near the cape in 1929 has occasionally been known as kap rink, but is officially known as nanok. (cape rink.) kap robert 72ø-419 (72°50.6´n 26°43.2´w). cape in ne glet scher land, named during the 1931–34 treårsekspeditionen by eugène wegmann as cape robert. kap russel [ilinnikajiip kiammut nuaa] 69ø-1 (69°58.7´n 22°24.6´w; map 4). cape sw of kap brewster. named in 1822 by william scoresby jr. as cape russel after james russell [1754– 1836], professor of clinical surgery at the university of edinburgh from 1803 to 1834. (c. rushel.) kap ruth 74ø-274 (74°04.8´n 22°17.0´w). cape on the se side of jordanhill at the front of wordie gletscher. the name seems first to have been used in the form kapp ruth on the 1932a nsiu map. girl’s name. kap ruth hytten 74ø (74°04.9´n 22°16.7´w). hut at the se point of jordanhill (kap ruth), built for the east greenland police authorities in 1938. it is now a ruin. kap ryder 69ø-17 (69°06.3´n 25°03.0´w; maps 3, 4). cape between barclay bugt and d’aunay bugt on the northern blosseville kyst. named by g.c. amdrup’s 1898–1900 expedition after carl hart vig ryder [1858–1923], leader of the 1891–92 expedition to east greenland. the cape was the south limit of ryder’s explorations, the point where the expedition met unbroken winter ice on 14 au gust 1892. kap schumacher 74ø-35 (74°40.2´n 20°08.0´w; map 4). peninsula south of kuhn ø opposite kap hamburg. named cap schumacher by karl koldewey’s 1869–70 expedition, after hermann albert schumacher [1839–1890], a jurist and historian, and one of the mem bers of the expedition organising committee (j. løve, personal communication 2010). he was later general consul in new york. the cape has also been called kapp agnes or kapp hynæs by norwe gian hunters. a norwegian hut 3 km west of the cape, occasionally referred to as kap schumacherhytten, is more usually known as holmslethuset. (kap schuhmacher.) kap seaforth 71ø-20 (71°47.3´n 22°49.6´w; map 4). headland on the west side of fleming fjord. named in 1822 by william scores by jr. in compliment to a much respected family of edinburgh. (cape seaforth, kap seaford, kap seagrave.) kap simpson 72ø-3 (72°08.1´n 22°11.6´w; map 3, 4). rounded se headland of traill ø. named in 1822 by william scoresby jr. as cape simpson. (kapp simpson.) kap simpson hytte 72ø (72°08.0´n 22°12.5´w). danish hut built by sirius in 1955–56 about 2 km west of kap simpson. it is also known as simpson-stranda. both name variations had also been given to an early proposed hut on this site, which was never built, although materials were deposited here by the veslekari in 1929. kap skt. jacques 77ø-4 (77°36.8´n 18°08.2´w; map 4). sw cape of île de france (from 2004 qeqertaq prins henrik), named by the duke of orléans in 1905 as cap st. jacques. eigil knuth suggested (personal communication, 1990) that it was named after rue saintjacques that runs through the area known as île de france in the centre of paris. (cap saint-jacques.) kap smith 71ø-4 (71°15.1´n 21°38.7´w; map 4). se point of tre kanten in eastern liverpool land. named cape smith by william scoresby jr. in 1822, together with other names in the vicinity for different friends, chiefly resident in manchester. possibly given for sir william sidney smith [1764–1840], an admiral in the royal navy who had corresponded with scoresby. kap steensby 76ø-41 (76°53.8´n 18°11.7´w; maps 2, 4). cape on the east coast of germania land. it was named during the 1906–08 danmark-ekspeditionen after hans peder steensby [1875–1920], who had assisted thostrup in preparation of his archaeological report. steensby was professor in geography at the university of copenhagen from 1911, and took part in several expeditions to africa, greenland and labrador. a hut built here by the norsk– franske polarekspedisjon in 1938, sometimes referred to as kap steensby hytten, has more usually been known as marga rine centralen. kap steglich-petersen 79ø-10 (79°31.0´n 22°19.3´w; map 4). cape-like feature in southern kronprins christian land facing south towards the inland ice. the area was mapped by lauge koch from the air in 1933 during the 1931–34 treårsekspeditionen, and the feature was named after the high-court lawyer kristian steglich-petersen [1880–1969]. he was denmark’s representative during the court action against norway at the hague concerning sovereignty of east greenland. kap stevenson [kangikajik] 70ø-73 (70°24.4´n 25°12.3´w; map 4). prominent headland 950 m high on the south coast of scoresby sund. named cape stevenson by william scoresby jr. in 1822, after robert stevenson [1772–1850]. an english civil engineer, steven son developed revolving lights, and designed and built 20 lighthouses, the most important being the bell rock lighthouse off the coast of angus, scotland. (kap stewenson, cape steven.) kap stewart [innakajik] 70ø-281 (70°26.6´n 22°38.2´w; maps 3, 4). se cape of jameson land. it was named cape stewart by wil liam scoresby jr. in 1822 after dugald stewart [1753–1828], who was professor of mathematics at the university of edinburgh from 1775, and of moral philosophy from 1785. the name was misspelt kap steward in the german edition of scoresby’s narrative (1825) and ejnar mikkelsen used the incorrect kap steward form in several of his descriptions of the scoresysund colonisation (e. mik kelsen 1950, 1989). the former settlement north of the cape was also occasionally called kap stewart, although was usually known as ittorisseq (see also tsuletsulekajik.) kap stop 76ø-113 (76°37.8´n 21°39.7´w; map 4). south cape of daniel bruun land, so named by j.p. koch’s 1912–13 expedition because their progress by boat was stopped here at the entrance of borgfjorden by dense glacier ice calved from bredebræ. the expedition waited until the fjord froze before continuing their journey by horse-drawn sledge. koch’s camp site, the skeletons of ponies, and a cairn were found here during the 1989 ggu expedition. (hindrunarhöfdi.) kap stophytten 76ø-201 (76°38.8´n 21°38.2´w). danish hunting hut on the north side of kap stop in southern daniel bruun land, built by nanok in september 1933, and replaced by a new hut in 1939. (kap stop hytten.) kap stosch 74ø-80 (74°03.6´n 21°43.8´w; maps 2, 4). north point of hold with hope. named by karl koldewey’s 1869–70 expedition as cap stosch for albrecht von stosch [1818–1895], a german general and admiral ( j. løve, personal communication 2010). norwegian hunters used kapp krogness for a minor cape near kap stosch, although it was often assumed to refer to the main cape (see kapp krogness). krogness was the name of the norwegian hunting station sw of kap stosch. (kapp stosch, cape stosch.) kap sussi 75ø-38 (75°19.1´n 17°47.9´w; map 4). cape on the east side of shannon, named by the 1909–12 alabama expedition as cape sussi. the name is unknown amongst present-day members of ejnar mikkelsen’s family, and may have been adopted from an unpublished chart by one of mikkelsen’s whaling associates. the remains of the german meteorological station of the 1943–44 operation bassgeiger are found nearby (75°19.2´n 17°48.1´n), together with the grave of lieutenant gerhard zacher shot here by 223 the sledge patrol on 22 april 1944 (fig. 1943–44 bassgeiger). see also fünkhütte. (cape suci). kap swainson [nuua] 70ø-335 (70°25.9´n 21°43.6´w; map 4). cape in southern liverpool land. named cape swainson by william scoresby jr. in 1822 in compliment to william swainson [1789–1855], a naturalist who made valuable zoological collections during travels to the mediterranean and brazil. a large hut has been built at the cape by scoresbysund municipality. (cape svainson). kap syenit 72ø-132 (72°03.4´n 23°06.3´w; map 4). cape on the nw side of antarctic havn, ne scoresby land. the name was proposed by lauge koch during the 1931–34 treårsekspeditionen, and first used by noe-nygaard (1934) in the form cape syenite. the cape is formed by a syenite intrusion. kap tattershall 71ø-3 (71°11.3´n 21°40.4´w; map 4). cape in ne liverpool land. named cape tattershall by william scoresby jr. in 1822, together with other features in the vicinity for different friends chiefly resident in manchester. kap thermopylæ 71ø-138 (71°04.4´n 21°54.4´w). cape on the south side of storefjord, so named during the 1931–34 treårs ekspeditionen by helge g. backlund because of the conspicuous hot springs. the locality thermopylae in east central greece is noted for its hot mineral springs. kap tobin 70ø-322 (70°24.9´n 21°58.0´w; maps 3, 4). radio and weather station at kap tobin [uunarteq], southernmost liverpool land. it was built in 1947 and closed down in 1980. an automatic weather station was erected in august 1985. some buildings were taken over by scoresbysund municipality for use by the kap tobin settlement, but most are now abandoned (see below). kap tobin [uunarteq] 70ø-323 (70°24.9´n 21°58.0´w; map 3). greenlandic village at kap tobin in southernmost liverpool land. the 1924–25 expedition that founded scoresbysund built two houses here. further houses were subsequently built, and the locality was permanently occupied until 2005. some of the weather station buildings abandoned after its closure in 1980 were taken over by the village, which had a population of 48 in 1990, but only six in 2000; there were no permanent residents after 2005. (pt. tobin.) kap tobin [uunartip nuua] 70ø-324 (70°24.6´n 21°56.7´w; map 3). southern cape of liverpool land, named by william scoresby jr. in 1822 as cape tobin, in compliment to sir john tobin [1763– 1851] of liverpool, merchant and ship-owner (fig. 3). the settlement near the cape has been known as kap tobin or uunarteq, and the radio station as kap tobin. scoresbysund town has occasionally used the name vardepynten for this cape. (cap tobin.) kap topham 71ø-6 (71°19.9´n 21°38.2´w; map 4). cape in north liverpool land, named cape topham by william scoresby jr. in 1822 after his friend john topham. kap toula 75ø (75°06.7´n 20°42.6´w). cape opposite kap negri at the mouth of grandjean fjord. the name is occasionally seen in reports by helge g. backlund on work during the 1931–34 treårs ekspeditionen (in: koch 1955). it was given for the austrian geologist franz toula [1845–1920], a contempory of christoforo negri (see kap negri), and well known for his studies of carboniferous faunas. kap tramnitz 75ø-30 (75°00.3´n 18°52.7´w; map 4). sw cape of shannon. named cap tramnitz by karl koldewey’s 1869–70 expedition after otto tramnitz [1847–1875], second officer of the expedition ship germania. he was drowned in a shipwreck in 1875. (cape tramnitz.) kap tramnitz hytten 75ø (75°03.9´n 18°54.0´w). danish hunting hut on the west coast of shannon, about 6 km north of kap tram nitz. it was built by nanok in september 1948, and is also known as tomsborg. (tramnitzhytten.) kap trekløver 77ø-44 (77°16.0´n 24°21.6´w; fig. 21). nw projection of prins axel nunatak, dronning louise land. named during the 1909–12 alabama expedition, probably by wilhelm laub, for its appearance (trekløver = clover). (cape tre kløver, kap treklöver.) kap tyrrell 71ø-16 (71°45.5´n 22°12.5´w; map 4; see also fig. 90). northern cape of canning land. named during the 1931–34 tre årsekspeditionen by arne noe-nygaard for george walther tyr rell [1883–1961], a british igneous petrologist noted especially for his work in scotland and his book on ‘the principles of petrology’. (cape tyrrel.) kap udkiggen 76ø-70 (76°43.6´n 18°26.1´w). south cape of ørnen ø, east of danmarkshavn. named by christian b. thostrup during the 1906–08 danmark-ekspeditionen for the danish petty officer association journal ‘udkiggen’ (j. løve, personal communication 2009). (cape look-out, kap udkiggeren.) kap ullidtz 76ø-46 (76°14.9´n 21°43.0´w; map 4). cape in rech nitzer land at the front of soranerbræen. so named by the 1906–08 danmark-ekspeditionen, possibly by henning bistrup after hans christian ullidtz [1878–1950], a captain in the danish navy. henning bistrup and h.c. ullidtz were promoted to second lieutenent on the same day (j. løve, personal communication 2009). a norwegian hunting hut built near the cape in august 1933 was known as sjelnan. kap ursus major 71ø (71°57.9´n 28°24.9´w). name used by helge g. backlund during the 1931–34 treårsekspeditionen for the east cape of charcot land (in: koch 1955), and given for the constellation. see ursus major gletscher. kap ursus minor 71ø (71°57.7´n 28°16.7´w). cape at the foot of backlund bjerg, inner nordvestfjord, so named during the 1931–34 treårsekspeditionen by helge g. backlund (in: koch 1955) after the constellation. see ursus minor gletscher. kap uttental 80ø (80°39.6´n 17°02.9´w). cape on the north side of ingolf fjord, named by elmar drastrup’s 1938–39 expedition after waldemar uttental, chairman of the scoresbysund committee that had supported the expedition. drastrup (1945) reported it as a cape immediately west of kap jungersen where he deposited a message in a cairn, whereas knuth (1942) reported this cairn to be at kap jungersen. kap vidar 71ø-113 (71°16.3´n 21°48.9´w). cape in eastern liverpool land west of trekanten. so named by helge g. backlund during the 1931–34 treårsekspeditionen, after his oldest son, vidar, who was his assistant in 1934. to avoid the recently introduced prohibition of naming features after living persons, back lund claimed it was named after the son of odin, god of norse mythology. kap wardlaw [ilittiartiip nuaa] 71ø-15 (71°44.2´n 21°54.1´w; map 4; see also fig. 90). ne cape of canning land, named cape wardlaw by william scoresby jr. in 1822 after robert wardlaw of tillicoultry. kap weber 73ø-502 (73°30.0´n 24°43.3´w; map 4). eastern cape of andrée land. named cap weber by karl koldewey’s 1869–70 expedition, possibly for wilhelm eduard weber [1804–1891], a german scientist who had worked with gauss. from 1849 he was head of göttingen observatory (j. løve, personal communication 2010). see also gauss halvø. (cape weber.) kap weinschenck 76ø-120 (76°58.9´n 23°09.9´w; map 4). low hill on the east side of dronning louise land, west of strandelv, named by j.p. koch’s 1912–13 expedition after ivar kjerulff weinschenck [1882–1963]. weinschenck was first engineer on the danmark during the 1906–08 danmark-ekspeditionen, and a chief engineer with the østasiatiske kompagni and other shipping companies. he had visited dronning louise land on a sledge journey in 1908. kap wijkander 73ø-28 (73°09.5´n 22°52.3´w; maps 3, 4). easternmost cape of ymer ø. named by a.g. nathorst’s 1899 expedition as wijkanders ö, probably after erik anders gustaf august wijkander [1849–1913], a swedish physicist and politician who had participated in the 1872–73 expedition to spitsbergen. in 1929 nsiu and lauge koch independantly made the observation 224 that the ‘island’ was connected to ymer ø by a low peninsula. (wijkander island, wijkander peninsula, c. wijkander, wijkander halvøya, wijkander-ø, kapp wijkander, kapp vikander.) kap wynn 74ø-17 (74°29.0´n 18°59.0´w; map 4). cape in eastern wollaston forland, named cape wynn by douglas clavering in 1823. several hunting huts were built about 1 km nw of the cape (see kopperneshuset, liavaag, gåsneshuset). (cap wynn, cape wyen, kap wyen.) kap young 72ø-5 (72°15.1´n 22°02.6´w; maps 3, 4). headland on se traill ø, named cape young by william scoresby jr. in 1822 after george young [1777–1848]. he became pastor of a presbyterian congregation at whitby in 1806, and stayed there 42 years. kap zachariae 79ø-37 (79°00.0´n 20°19.0´w; map 4). peninsula on the south side of lambert land, north of zachariae isstrøm. named by john haller following explorations during lauge koch’s 1956–58 expeditions. kap øtker – see kap oetker. kap øtker hytten – see nes-odden. kap aage bertelsen 76ø-122 (76°40.1´n 23°03.0´w; map 4). minor feature in eastern dronning louise land forming a small cape-like feature at the confluence of storstrømmen and l. bistrup bræ. named by j.p. koch’s 1912–13 expedition after aage ber telsen [1873–1945], artist on the 1906–08 danmark-ekspedi tionen. bertelsen and achton friis made several hundred paintings and drawings during the expedition. kapelle 72ø (72°01.1´n 25°10.1´w; map 5). mountain on the ne side of sefström gletscher, stauning alper. named and first climbed by hans gsellman’s 1957 expedition. kapelleturm 72ø (72°01.5´n 25°09.2´w). name used by hans gsellman’s 1957 expedition for the present beaufort tinde, stau ning alper, located on the ne side of sefström gletscher ne of kapelle. their attempt on the peak was frustrated, and it was first climbed in 1958 by malcolm slesser’s party. kaphytten 75ø (75°56.3´n 19°57.8´w). norwegian hunting hut at kap møbius, south of the mouth of bessel fjord, built by john giæver’s expedition in november 1932. kaporniagaqarpik [konglomeratelv] 71ø-226 (71°20.2´n 24°48.7´w). river draining from the eastern lake of holger danske briller into the west side of nordøstbugt, east of sydkap. recorded by the 1955 geodætisk institut name registration, the name translates as ‘where there are trout’. kaporniagaqarpik 71ø-204 (71°26.6´n 25°19.7´w). river draining westwards from the west lake of holger danske briller, southernmost stauning alper. recorded by the 1955 geodætisk institut name registration, the name means ‘where there are trout’. kaporniagaqarteq [søelv] 70ø-166 (70°43.9´n 22°24.2´w). river on the east side of hurry inlet draining sødal. recorded during the 1955 geodætisk institut name registration, the name translates as ‘it has trout’. kapp – see also cap, cape and kap. kapp 17. mai 72ø (72°53.5´n 24°31.6´w). cape on western geo graphical society ø, so named on the nsiu maps of lacmann (1937) for norway’s national day. (cape 17th of may.) kapp 7. juni 72ø (72°58.9´n 24°33.5´w). cape on west geographical society ø. so named on the nsiu maps of lacmann (1937) after the 7de juni, a 14-ton, 40-foot sealer used by the pioneer 1909–10 wintering expedition led by vebjørn landmark. kapp agnes 74ø (c. 74°40´n 20°14´w). norwegian hunting hut on the south side of lindemann fjord, near kap schumacher, built by the hird expedition in september 1928. the norwegians often referred to kap schumacher as kapp agnes or kapp hynes. agnes was the youngest daughter of jørgen furnes, who helped move the hut to this site from kap stosch. she was born after furnes left for greenland in 1927. the hut was moved in august 1930 to kap hamburg on kuhn ø. it has also been known as furnes. (hyneshytten, agnes-tufta.) kapp astrid 74ø (74°19.2´n 22°03.2´w). minor cape in southern payer land on the north side of the mouth of grantafjord. this position is shown on the 1:100 000 scale nsiu maps (lacmann 1937), but on the 1932a nsiu map it appears to be indicated as the cape on the opposite side of the fjord, the present grantapynt. the latter usage was adopted by den grønlandske lods (1968). (kap astrid.) kapp bjørvig 74ø (74°26.2´n 20°56.2´w). cape on the west side of lerbugt, north clavering ø. so named on the nsiu maps of lacmann (1937) after paul bjørvig [1857–1932], a tromsø hunter who had participated in expeditions to the arctic and antarctic. kapp blosseville – see blosseville bjerg. kapp brandal 74ø (74°24.9´n 21°37.0´w). cape on nw clavering ø. used only on nsiu maps (lacmann 1937), the name commemorates peter severinsen brandal [1870–1933] of brandal, the norwegian ship-owner who instituted norwegian sealing activities off east greenland. kapp cathrine 74ø (74°18.2´n 22°12.0´w). cape on the south side of grantafjord. used on the nsiu maps of lacmann (1937). girl’s name. kapp dagny 74ø (74°18.0´n 22°20.4´w). cape in southernmost payer land on the north side of grantafjord. used only on the nsiu maps of lacmann (1937). kapp else 74ø (74°05.0´n 22°24.2´w). minor cape on the south side of jordanhill. so named on the nsiu maps of lacmann (1937). kapp floren 73ø (73°00.2´n 24°11.0´w). minor cape on the north side of western geographical society ø. so named on the nsiu maps of lacmann (1937), after the floren, the sealer used by the 1908–09 floren expedition led by severin liavaag. the 37-ton floren was constructed by hans gravdal of opsanger, and was the first ship built in sunnmøre for arctic use. see also kapp liavåg. kapp giæver 74ø (74°11.4´n 22°13.2´w). minor cape north of jordanhill. so named on the nsiu maps of lacmann (1937) after john giæver. see also giæverhytte. kapp gjöa 72ø (72°54´n 24°17´w). minor cape on the south side of geographical society ø. the name is used only on nsiu maps (lacmann 1937), and was given for the gjöa, the 47-ton herring boat with which roald amundsen made his voyage through the nw passage in 1903–05. it is now a museum ship in oslo. kapp grödahl 74ø (74°17.5´n 20°25.8´w). cape on ne clavering ø, corresponding to the delta of storstrømmen. used only on nsiu maps (lacmann 1937), the name was given for ole iversen grødahl [1850–1922], a norwegian skipper who pioneered summer sealing off east greenland. kapp hedlund hytta 72ø (72°43.1´n 26°10.5´w). norwegian hunting hut in the bay east of kap hedlund, built by arktisk næringsdrift in 1934, and also known as rimhytten. it was replaced in 1964 by a new hut built by sirius. the skeleton of an unnamed hut occurs on the west side of kap hedlund; strong winds from rhedin fjord prevented its completion, and kap hedlund hytta (rimhytten) was built instead (p.s. mikkelsen 1994, 2008). (kapp hedlund.) kapp hekla 72ø (72°56.0´n 24°34.5´w). minor cape on west geographical society ø. so named on the nsiu maps of lacmann (1937) after the norwegian ship hekla. see hekla havn. kapp herschell – see herschellhus. kapp hynæs – see kapp agnes. kapp isachsen 73ø (73°13.2´n 23°16.3´w). cape on the north side of the mouth of dusén fjord, sw of kap graah. named on an nsiu map (1932a) after gunnerius ingvald isachsen [1868–1939], a nor we gian polar explorer who led several expeditions to spits bergen, and the 1930–31 ‘norvegia’ expedition to the antarctic. kapp isbjørn 72ø (72°51.2´n 23°01.2´w). minor cape on the south side of geographical society ø. so named on nsiu maps of lacmann (1937) after the isbjørn, a 172-ton norwegian sealer built in 1918, and used by a variety of norwegian and foreign expe225 ditions for voyages to franz josef land, svalbard and greenland. kapp johan olsen 74ø (74°15.3´n 21°59.8´w). cape on west clave ring ø, the present kap oetker. the name is used only on nsiu maps (lacmann 1937), and was given for johan peter kornelius olsen [b. 1879] who as skipper of the veslekari made great contributions to the scientific expeditions of nsiu in east greenland. he is said to have found the cod banks off west greenland and to have opened up the fishery in 1925. kapp krogness 74ø (74°02.8´n 21°46.8´w). minor cape close to the norwegian hunting station krogness, sw of kap stosch. named by the 1926–28 foldvik expedition after ole andreas krogness (see also krogness) who had given them great help and advice, and stimulated them to undertake the expedition. this was the first place where the expedition landed. the hunting station has also often been referred to as kapp krogness. for many years it was assumed that kapp krogness was the norwegian name for kap stosch, but this was a misunderstanding (svend bendix-almgren, personal communication 1997). kapp landmark 74ø (74°07.0´n 20°46.7´w). cape on the se side of clavering ø. so named on the nsiu maps of lacmann (1937) after vebjørn landmark [b. 1879], who led the hunting expedition which overwintered in east greenland in 1909–10 with the 7de juni. he was mate on the veslekari in 1929 and the polarbjørn in 1930, during nsiu expeditions to east greenland. the cape has also been called cap alf. kapp laura 72ø (72°52.4´n 23°26.1´w). minor cape on the south side of central geographical society ø. used on the nsiu maps of lacmann (1937), it was named after the norwegian sealer laura. kapp liavåg 74ø (74°14.1´n 20°18.2´w). name used for the delta on east clavering ø at the mouth of grønnedal on the nsiu maps of lacmann (1937). it was named after severin gaasnes liavaag [1879–1909], who was leader of the 1908–09 hunting expedition to the region. see also gåsneshuset. kapp lillenæs 74ø (74°12.0´n 22°11.3´w). minor cape north of jordanhill. used on the nsiu maps of lacmann (1937), the name commemorates paul lillenæs [b. 1877], skipper of the veslekari which carried the nsiu expedition to east greenland in 1930, and louise boyd’s expedition to the same region in 1931. kapp martha 73ø (73°19.0´n 23°31.4´w). cape on the ne side of of ymer ø, so named on an nsiu map (1932a). a hunting hut at the cape sometimes known as kapp martha hytten is better known as slippenhytten. kapp marö 74ø (74°24.3´n 21°47.9´w). name used for the delta on the east coast of payer land south of kap ehrenberg on nsiu maps (lacmann 1937). the name was given for kristoffer marø [b. 1884], skipper of the polarbjørn which was extensively used by nsiu expeditions to east greenland. during the 1939–45 war marø with the polarbjørn carried ammunition and supplies to arctic waters for the united states, and the polarbjørn acquired the reputation of ‘the ship that always arrives’. kapp minerva 72ø (72°52.1´n 23°14.0´w). minor cape on the south side of central geographical society ø. so named on the nsiu maps of lacmann (1937) after the norwegian sealer minerva of tromsø, occasionally used to carry expeditions to greenland. (kap minerva). kapp minna 72ø (72°54.5´n 24°00.0´w). minor cape on the south side of west geographical society ø. so named on the nsiu maps of lacmann (1937) for the minna, a 68-ton norwegian sealer built in hardanger in 1894, that under the command of peter s. brandal initiated norwegian sealing off the coast of east greenland. kapp myklebust 72ø (72°46.7´n 22°57.4´w). north cape of kista ø in vega sund. used only on nsiu maps (lacmann 1937), it was named after johannes myklebust [b. 1894], who visited east greenland as skipper of the buskø in 1935. kapp norge 74ø (74°42.4´n 20°03.8´w). name sometimes used by norwegian hunters for kap hamburg in southern kuhn ø, which they also called røsnes. kapp norge has also been used for the norwegian hunting hut west of the cape, usually known as furnes. kapp næssø 74ø (74°23.2´n 21°43.1´w). cape on nw clavering ø. used only on nsiu maps (lacmann 1937), and named after ole næs sø [1844–1921], a norwegian skipper who made many sum mer hunting expeditions to east greenland. kapp oetker 74ø (c. 74°15´n 22°00´w). norwegian hunting hut at kap oetker, west clavering ø, built in august 1927 by the foldvik expedition. it was moved in 1929 to eskimovig. kapp petersens – see kap peterséns. kapp polarbjørn 73ø (73°03.8´n 23°13.3´w). minor cape on the north side of central geographical society ø, west of robertson ø. so named on the nsiu maps of lacmann (1937) after the polar bjørn, a 360-ton sealer built in 1919 and used as an expedition ship by nsiu and arktisk næringsdrift from 1932 to 1939 and from 1946 to 1948. it was lost by fire off newfoundland in 1949. in the war years, with kristoffer marø as skipper, it was used by the united states for transporting ammunition and supplies to the arctic (see kapp marø.) kapp quest 72ø (72°59.3´n 24°26.0´w). minor delta on the north side of geographical society ø. so named on the nsiu maps of lacmann (1937) for the norwegian sealer quest. built as a sealer in 1917, it went under the name foca 1 until its purchase for the shackleton-rowett antarctic expedition of 1921–22. the quest was subsequently used for a number of arctic expeditions. it picked up the teddy expedition crew from ammassalik in 1924, brought home umberto nobile after his failed attempt to reach the north pole in 1928, and transported the british arctic air route expedition to east greenland in 1930. in 1962 it was lost in the ice off labrador. kapp ragnvald knudsen 74ø (74°24.3´n 20°33.2´w). cape on ne clavering ø, the delta at the mouth of dolomitdal. used only on nsiu maps (lacmann 1937), it was named after ragnvald knud sen [1858–1930], who as skipper of the hekla made one of the earliest norwegian hunting visits to east greenland in 1889. with the hekla he sailed carl ryder’s 1891–92 expedition to the scores by sund fjord complex, and is said to have discovered two new fjords on the blosseville kyst. kapp randi 74ø (74°19.1´n 22°05.3´w) cape on the north side of the mouth of grantafjord. the name is only found on the nsiu maps of lacmann (1937). kapp ringsæl 74ø (74°54´n 23°48´w). minor cape on the south side of west geographical society ø. so named on the nsiu maps of lacmann (1937) for the ringsel of tromsø, a norwegian sealer which made several visits to east greenland. in 1938–39 it was renamed en avant for the duration of gaston micard’s norsk– franske polarekspedisjon. it was lost off east greenland in 1952. photographs of the sealer clearly show the spelling ‘ringsel’. kapp rygg 72ø (72°51.8´n 23°33.7´w). norwegian hunting hut on the north side of vega sund, east of the pronounced cape marked on norwegian maps as kapp rygh. the hut was built by arktisk næringsdrift in 1929. it was also known as kapp rygh, rev-odden, solveigs hytten and sverdrup hytte. kapp rygh 72ø (72°51.8´n 23°35.0´w). pronounced cape on the south side of geographical society ø. so named on the nsiu maps of lacmann (1937) after oluf rygh [1833–99], a norwegian archae o logist and historian. the name was also adopted for the norwegian hunting hut east of the cape (see kapp rygg). kapp sandefjord 74ø (74°26.6´n 20°25.9´w). cape on the sw coast of wollaston forland. the name is used only on nsiu maps (lac mann 1937), and was given for the district of sandefjord in nor way, the home of several important whaling companies. kapp schjelderup 74ø (74°18.8´n 21°55.3´w). cape on west clave ring ø, the delta at the mouth of tørelv. so named on the nsiu maps of lacmann (1937) after ludolf schjelderup [b. 1894], a noted norwegian sealer skipper. he captained the quest during 226 expeditions to east greenland and svalbard. kapp sjøblomsten 73ø (73°00.5´n 23°53.5´w). minor cape and delta on the north side of west geographical society ø. so named on the nsiu maps of lacmann (1937) for the sjøblomsten, a norwegian sealer which visited east greenland in 1912. (kapp sjöblomsten.) kapp sulabak 73ø (73°53.4´n 20°01.9´w). se cape of jackson ø, named in this form on an nsiu map (1932a). named after peder sulabak, a member of the 1927–29 hird expedition which operated in this area. he was also a member of the 1930–32 møre expedition. (cape sulabak.) kapp sælbarden 73ø (73°01.5´n 23°38.9´w). minor cape and delta on the north side of central geographical society ø. named after the norwegian sealer sælbarden of ålesund, used by nsiu expeditions in 1934. it was wrecked in 1937. kapp thor iversen 72ø (72°38.8´n 22°42.6´w). cape on the ne side of traill ø, west of nordenskiöld ø. used only on nsiu maps (lacmann 1937), the name commemorates thor iversen [1873– 1953], leader for many years of the fiskeri direktoratet (directo rate of fisheries) in bergen and responsible for dispatch of numerous expeditions to arctic waters. kapp tromsø 73ø (73°59.2´n 21°59.4´w). minor spit on the large delta on the west side of loch fyne. the name is used on the nsiu maps of lacmann (1937), and was given for the town of tromsø, the traditional departure point of norwegian arctic expeditions. (tromsöyra.) kapp veslekari 73ø (73°02.6´n 23°28.2´w). minor cape and delta on the north side of geographical society ø. so named on the nsiu maps of lacmann (1937) after the veslekari, a 282-ton, 125-foot sealer built in 1918 for svend foyn and extensively used for sealing in spitsbergen, greenland and newfoundland waters. it was often used as an expedition ship to east greenland, in 1929 and 1930 with nsiu expeditions, and in 1931, 1933, 1937 and 1938 with louise boyd’s expeditions (ellefsen & berset 1957). it was still considered one of norway’s best sealers when lost off newfound land in 1960. kapp wollebæk 72ø (72°50.1´n 23°10.0´w). cape on the north side of central traill ø, the present østernæs. so named on the nsiu maps of lacmann (1937) after alf wollebæk [1879–1960], a nor we gian zoologist who became director of the zoological museum in oslo. veganeset has been used on norwegian maps for the same feature. kapp øien 74ø (74°08.7´n 21°30.0´w). cape on sw clavering ø, equivalent to the delta at the mouth of granatdal. the name is used on the nsiu maps of lacmann (1937), and was given for jens øien [b. 1870], a norwegian skipper who with the laura sailed a number of hunting expeditions to east greenland. kapp ålesund 74ø (74°07.3´n 22°10.6´w). cape on the east coast of jordanhill. used on the nsiu maps of lacmann (1937), the name was given after the town of ålesund in norway, home of many of the norwegian sealers that hunted off east greenland. kapspidsen 76ø-10a (76°12.5´n 19°57.0´w; map 4). mountain near kap peschel in ad. s. jensen land. the name kap-spitze is only mentioned in the geology section of karl koldewey’s 1869–70 expedition narrative, but was adopted by subsequent visitors to the region and approved in its danish form. kaptajn hansens promenade 74ø (74°42.7´n 18°16.0´w). name given by danish hunters to a pathway constructed by captain f. hansen on bass rock to improve the passage from the beach to the higher parts of the island. after the wreck of the dagny commanded by hansen in 1920, the nine crew and danish hunters wintered at bass rock and shannon. kar glacier 74ø (74°29.8´n 19°18.4´w). name used by andreas vischer (in: koch 1955) in a report on his 1937 field work, for a glacier on the slopes of hühnerbjerg east of point 630 m (kar = large vessel or bathtub). karabiner fjeld 71ø-341 (71°37.5´n 24°57.0´w; map 5). mountain 2000 m high south of leo gletscher, southern stauning alper. first climbed by john hunt’s 1960 expedition, and named karabiner for the karabiner mountaineering club of which he was honorary president. the second ascent was by the 1971 university of lan caster expedition. karbon elv 74ø (74°24.8´n 20°15.7´w). river flowing through sandstensdal, west wollaston forland. the name was used by alfred rosenkrantz (1932) because rocks of carboniferous (= kar bon) age were found here in 1929. (karbon river.) karboncircus bjerg – see circusbjerg. kargletscher 71ø-267 (71°58.2´n 24°01.4´w; map 5). small glacier in the werner bjerge, merging to the north with østre gletscher. named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk. karhøjen 71ø-268 (71°57.9´n 23°58.5´w; map 5). mountain in the werner bjerge between kargletscher and østre gletscher, named by peter bearth and eduard wenk during the 1953–54 lauge koch expeditions. it was climbed by bearth in 1953. karin dal 73ø-74 (73°30.7´n 22°47.8´w). valley on gauss halvø draining north into moskusoksefjord, named by lauge koch’s 1929–30 expeditions in the form karin valley. girl’s name, said to be a swedish girlfriend of one of the expedition members. (karins dal, karin tal.) karina 74ø (74°18.4´n 20°13.6). wintering house at sand odden/ daneborg, said to have been built by the scoresbysund committee about 1938. the present hotel karina at daneborg has been con verted to a museum to trapping activities (p.s. mikkelsen 2008). karinas lyst 70ø (70°29.1´n 21°57.9´w). name given to the first small house built by the 1924 expedition that founded scoresby sund; it was a food store. it was named after karina bell, a danish actress who was aage nielsen’s cousin. see also aage nielsen bjerg. karl dal 73ø-342 (73°32.5´n 22°04.8´w). valley in the northern giesecke bjerge draining east into badland dal. named during lauge koch’s 1936–38 expeditions by wolf maync and andreas vischer, after karl andersen, their greenlandic assistant and sledge-driver in 1937 and 1938. karl jakobsen bugt 73ø-558 (73°03.3´n 24°44.0´w; map 4). bay on the south coast of ymer ø, named by j.m. wordie’s 1929 expedition as karl jakobsen bay after the skipper of the heimland which carried the expedition to greenland. a norwegian hut on the coast of the bay sometimes known under the name karl jakobsen bugt is better known as namdalsstua. (k. jakobsens bugt.) karl pynt 75ø-58 (75°14.6´n 20°01.2´w; map 4). peninsula on the south side of lauge koch vig, southern hochstetter forland. named by hans frebold during the 1931–34 treårsekspeditionen. (karls pynt.) karlenes ø 72ø-100 (72°26.7´n 24°46.0´w; map 5). island at the mouth of segelsällskapet fjord. named during the 1931–34 tre årsekspeditionen by ove simonsen in tribute to the crew of a.g. nathorst’s 1899 expedition ship, frequently referred to in the expedition narrative as ‘karlarne’ (= the crew). (karlenes insel.) karlsbak 71ø (71°59.7´n 23°06.7´w). hunting station in the inner part of antarctic havn, erected for the møre expedition in august 1930 by jonas karlsbak and odd åmbak. it was manned in the periods 1930–38 and 1946–59. the station has also been known under the names bakkehuset, antarctic havn station and antarc tichamna. it was restored by nanok in the summer of 2001, but destroyed in an avalanche the following winter. karlsbakfjellet 74ø (74°08.7´n 20°51.0´w). mountain on south clavering ø, the south ridge of the present pladen. used only on nsiu maps (lacmann 1937), the name was given for the norwe gian hunter jonas karlsbak [b. 1895], who wintered in east green land in 1927–29 and 1930–31. karlshavn – see carlshavn. karstgraven 71ø-307 (71°28.3´n 24°33.2´w). valley in the south part of the the karstryggen area, which shows characteristic karst 227 features. named by enrico kempter during lauge koch’s 1956–58 expeditions (graven = the grave). karstryggen 71ø-158 (73°30.0´n 24°37.8´w). ridge west of schuchert dal in which a thick dolomite bed gives rise to karst topography. named by hans stauber during lauge koch’s 1936–38 expeditions. karupelv 72ø-89 (72°32.6´n 23°43.1´w; map 4). river on sw traill ø, named by ove simonsen during the 1931–34 treårsekspedi tionen after the danish river karup å in jylland. karupelv hytten 72ø (72°30.1´n 24°00.3´w). name sometimes used for the norwegian hut built in july 1932 at the mouth of karupelv in holm bugt, traill ø. it was restored by nanok in 2001. see also holm-vika. kaskadesø 70ø-377 (70°15.4´n 28°58.1´w). lake in west gåseland surrounded by waterfalls (= kaskade). named during lauge koch’s 1958 expedition by eduard wenk. the pilots of the catalina that landed wenk’s party here called it blå sø. kassen 71ø-398 (71°35.5´n 22°53.2´w). mountain 942 m high on sw wegener halvø. named by katherina perch-nielsen during the 1967–72 ggu scoresby sund expeditions, for its angular shape (kasse = box). kastellet 70ø-360 (70°08.4´n 22°11.3´w). mountain 441 m high west of kap brewster, savoia halvø, named during the 1931–34 treårsekspeditionen by laurits bruhn for its appearance (kastellet = the citadel). kastenberg 71ø (71°59.0´n 25°12.5´w; map 5). mountain on the sw side of sefström gletscher, stauning alper. first climbed by hans gsellman’s 1957 expedition. katederet 71ø-273 (71°54.2´n 24°13.0´w). mountain between arcturus gletscher and sirius gletscher. named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk (katederet = the pulpit). katederryggen 71ø-274 (71°56´n 24°15´w; map 5). ridge between arcturus gletscher and sirius gletscher, west werner bjerge. named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk (see katederet). katederspidsen 71ø-275 (71°56.1´n 24°12.9´w; map 5). mountain between arcturus gletscher and sirius gletscher, west werner bjerge. named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk (see katederet). katedralen 70ø-112 (70°52.8´n 22°57.1´w). mountain 610 m high in eastern jameson land west of the head of hurry inlet, named by alfred rosenkrantz during lauge koch’s 1926–27 expeditions in the form cathedral mt, after its shape. kater bay 74ø (74°31.5´n 19°05.8´w). this is probably identical with claveringstrædet between wollaston forland and sabine ø. william scoresby jr. in 1822 had named kater bay after henry kater [1777–1835], for many years treasurer of the royal society, and noted for his pendulum experiments. scoresby observed his kater bay from a great distance, and it could not be definitely located by subsequent visitors. (kater’s bay). kathedrale 72ø (72°00.2´n 25°19.5´w). alternative name used in a report of hans gsellman’s 1957 expedition (koglbauer 1965) for attilaborgen opposite their main camp on the upper reaches of sefström gletscher, stauning alper (kathedrale = cathedral). käthispids – see after kærelv (‘ä’ is treated as ‘æ’ in danish). katinkakkut nunat 70ø-365 (70°29´n 21°58´w). cape on the west side of scoresbysund. recorded by the 1955 geodætisk institut name registration, the name means ‘katinka’s land’. digby & digby (1954) record that katinka, a resident of scoresbysund, was held in respect and awe by everybody, and held territorial rights over the rocky point where her house was built and the adjacent waters. (katinkákut nûat). katinkákut nûat – see katinkakkut nunat. kavalerfjorden 76ø-153 (76°32´n 22°00´w; map 4). narrow fjord which almost divides lindhard ø. named by j.p. koch’s 1912–13 expedition as kavaller fjorden, after one of the expedition’s icelandic ponies (kavalleren) which became stuck in loose snow here in march 1913. kayak vig 71ø (71°19.1´n 24°49.8´w). small bay on the west side of nordostbugt where the river draining holger danske briller enters the sea. the name was used by hall (1966) in his description of birds observed during the 1962 oxford university expedition. it was considered a suitable place to bring small boats ashore. keferstein 74ø-53 (74°37.2´n 18°59.9´w). mountain 699 m high on sabine ø. named by karl koldewey’s 1869–70 expedition as kefersteinberg, probably after wilhelm moritz keferstein [1833– 1870], professor of zoology at göttingen (j. løve, personal com munication 2010). (mt keferstein). kegle i 71ø-81 (71°43.8´n 22°38.1´w). cone-shaped mountain east of tvekegledal, wegener halvø, named during the 1931–34 tre årsekspeditionen by arne noe-nygaard as conus i. kegle ii 71ø-82 (71°43.6´n 22°37.5´w). cone-shaped mountain east of tvekegledal, wegener halvø, named during the 1931–34 treårsekspeditionen by arne noe-nygaard as conus ii. keglebjerg 74ø-184 (74°31´n 23°19´w). mountain about 1450 m high on the north side of wordie gletscher. the mountain was climbed by th. johansen and curt teichert on 23 march 1932 in the course of a journey along the inland ice margin during the 1931–34 treårsekspeditionen. the name was given by johansen, and used first by teichert (1933) and gelting (1934). in their original map reproduced in koch (1940; fig. 34) the name kentebjerg is used. both names refer to its cone-like shape. kegleformet top 73ø (c. 73°24´n 23°07´w). mountain on southern gauss halvø with a cone-like shape, possibly one of the hjelm bjergene on southern gauss halvø. the name appears on one of the folding maps of carl ryder’s 1891–92 expedition. keglen 80ø-63 (80°24.6´n 21°08.6´w; map 4). mountain 949 m high on the west side of southern vandredalen, south of port fjeldet, named by elmar drastrup’s 1938–39 expedition. this cone-shaped mountain was used as a surveying mark, and its position is clearly shown on eigil nielsen’s (1941) and drastrup’s (1945) maps. the 1957 ams maps place the name against the higher flat-topped mountain to the ne known as brockmeyer bjerg. keglerne 71ø-82a (71°43.7´n 22°37.9´w). common official name for kegle i and kegle ii, two cone-shaped mountains east of tve kegledal on wegener halvø. so named by arne noe-nygaard during the 1931–34 treårsekspeditionen. kehlers havn 70ø (70°26.9´n 26°14.7´w). helge vedel’s diaries of carl ryder’s 1891–92 expedition (gulløv 1991) indicate that this was the name originally used for the present hekla havn, southern danmark ø. kejser franz joseph fjord 73ø-17 (73°15´n 22°50´w – 73°08'n 27°44´w; maps 3, 4; see also fig. 68). major e–w-trending fjord system, bounded by suess land and ymer ø to the south, and frænkel land, andrée land and gauss halvø to the north. it was discovered and partially explored by karl koldewey’s 1869–70 expedition and named kaiser franz josephs fjord, after franz joseph karl von habsburg [1830–1916], emperor of austria from 1867. he made substantial donations to the expedition finances. norwegian maps use the spelling franz josef fjord. (kaiser franz josef fjord, kejsar frans josefs fjord, frans josefs fjord, frans josefs inlet, frantz joseph fjord, emperor franz-joseph’s fjord, fiord françois-joseph, le fjord de l’empereur franz joseph.) kelhofer gletscher 73ø-722 (73°10.0´n 26°24.9´w). glacier in suess land draining nw from payer tinde. the name commemorates a swiss naturalist, ernst kelhofer [1877–1917], and was said to have been suggested in the 1930s by swiss geologists for the present sonklargletscher. eugène wegmann and heinrich bütler were both students of kelhofer. the name was revived in 1969 at the suggestion of kelhofer’s daughter, but relocated to a glacier 15 km west of sonklargletscher. 228 kelvin klippe 76ø-311 (76°57.9´n 24°55.8´w; map 4). cliff south of admiralty gletscher in dronning louise land. one of the names given by the 1952–54 british north greenland expedition for notable scientists, it commemorates the scottish physicist lord kelvin [1824–1907]. he was professor of natural philosophy at the university of glasgow from 1846, and was particularly noted for his role in the development of the conservation law of energy and the absolute temperature scale. kempe fjord 72ø-43 (72°48.0´n 25°50.0´w; maps 3, 4; fig. 52). wide e–w-trending fjord between suess land and lyell land. named by a.g. nathorst’s 1899 expedition after the most generous supporter of the expedition, seth michael kempe [1857– 1946], a successful stockholm businessman. he was a good friend of per dusén, surveyor on the expedition. (kempes fjord, kempe fiord, kempefjorden, kjempefjorden.) kemptner horn 71ø (71°48.5´n 25°06.4´w; map 5). mountain 2337 m high on the ridge between roslin gletscher and mars gletscher. climbed by karl herligkoffer’s 1966 expedition, and possibly named after the bavarian town of kempten. kensington 72ø (72°08.5´n 24°52.6´w). mountain 2600 m high at the head of bersærkerbræ and skoldungebræ, north stauning alper, the present pyramidefjeld. first climbed by the 1963 im perial college expedition, and named after the royal borough of kensington in sw london, merged with chelsea in 1965. the se cond ascent was made by toni gobbi’s 1967 party. kentebjerg – see keglebjerg. ker doumer 70ø (c. 70°30´n 21°57´w). name of the 1932–33 french international polar year station at scoresbysund, south liverpool land, which was named after paul doumer [1857–1932], a friend and supporter of jean-baptiste charcot who helped establish the station. doumer was president of france when assassinated in 1932. nyholm-poulsen (1985) described the station in 1933 as comprising two buildings connected by a long passage. the building was subsequently used as a telegraphists house, and later as a hospital. a new hospital was built in 1957. see also doumer høj. (station poul doumer.) ker virginie 70ø (70°31.3´n 21°53.3´w). name used for a house erected in south liverpool land for the french international polar year 1932–33. it was apparently built on a 425 m high col ne of scoresbysund by the crews of the french ships pourquoi pas? and pollux, and named after virginie hériot [1890–1932], a french sailor who won a gold medal in the 1928 olympics. she is said to have made generous contributions towards the expenses of the expedition. kerberus 73ø-709 (73°11.5´n 28°33.5´w; map 4). nunatak 2500 m high north of petermann bjerg, western frænkel land. so named by john haller and eduard wenk following explorations during lauge koch’s 1951 expedition, because it resembled in shape a sitting dog. kerberus in greek mythology was a dog which guarded the entrance to a tomb. haller’s party climbed the mountain on 15 august 1951. kerstin dal 73ø-283 (73°28.8´n 23°15.9´w). small valley on gauss halvø draining north to paralleldal. named during the 1931–34 treårsekspeditionen by gunnar säve-söderbergh after his wife britta kerstin arnell [d. 1952]. (kerstin valley.) keswicktinde 71ø (71°57.5´n 25°05.5´w; map 5). peak 2430 m high in the stauning alper. climbed by the 2001 scottish mountain eering club expedition. ketilfjellet 73ø (73°23.9´n 23°04.0´w). mountain 1502 m high on the south side of gauss halvø, corresponding to the present nat horst bjerg. so named on an nsiu map (1932a), after ketil, one of the original norse settlers of greenland. (mt. ketil.) kiammut nuukajia [kap hodgson] 70ø-206 (70°33.5´n 21°30.3´w). north-facing cape in se liverpool land. one of the names recorded by the 1955 geodætisk institut name registration, the name translates as ‘its small cape in the north’. (kiámut nûkajia, kiammut nuukajia.) kiámut nûkajia – see kiammut nuukajia. kiderlen kløft 70ø-40 (70°43.7´n 25°17.9´w). ravine on the east coast of milne land between charcot havn and kap leslie. named by hermann aldinger during the 1931–34 treårsekspedi tionen as kiderlenschlucht or kiderlen schlucht. (kiderlens kløft, kiderlen ravine.) kieferer toppen 71ø (71°52.1´n 25°20.5´w; map 5). mountain 2430 m high on the south side of the head of roslin gletscher. climbed by karl herligkoffer’s 1966 expedition on 17 august, and named after the small town of kiefersfelden in the bavarian alps, birthplace of gebhart plangger, one of the climbers. (kieferner toppen.) kigkaqángitseq – see kikkaqanngitseq. kikiakajiip qaqqartivartaa 70ø-347 (70°05´n 22°28´w). summits on savoia halvø, on the south side of scoresby sund. recorded by the 1955 geodætisk institut name registration, the name means ‘kikiakajiip’s big mountains’. (kikiakajîp qáqartivartâ.) kikiakajik 70ø-348 (70°03.1´n 22°17.3´w). valley or ravine on the se side of savoia halvø. one of the names recorded by the 1955 geodætisk institut name registration, it means ‘the little ravine’. kikiakajîp qáqartivartâ – see kikiakajiip qaqqartivartaa. kikkaqanngitseq 70ø-230 (70°43.5´n 21°43.3´w). cliff on the south side of sandbach halvø, south liverpool land. one of the names recorded by the 1955 geodætisk institut name registration, it translates roughly as ‘it has no pinnacles’. (kigkaqángitseq.) kikut 73ø (73°10.6´n 23°08.3´w). norwegian hunting hut on the coast of ymer ø, south of the vinterøer, built by arktisk nærings drift in august 1929. the name is a norwegian expression for a locality with a good view. the hut is also known as dusens fjord hytten, and occasionally as steffensen. kildedal 75ø-39 (75°15.4´n 21°03.7´w; map 4). valley on the south side of ardencaple fjord. so named for the warm springs (= kilde) discovered here by the danish hunter andreas hvidberg in 1931. the valley was known at the time as blaabærdal. large, clear ice-domes develop above the springs in the winter, but the water temperature is said to be only a few degrees above freezing so that the springs are not conspicuous in the summer. (kildedalen.) kildedalhytten 75ø-103 (75°15.7´n 20°54.4´w). danish hunting hut on the north side of the mouth of kildedal, ardencaple fjord, built by nanok in september 1931. now a ruin (1988). (kildedal hytten.) kildeelv 74ø (74°27.9´n 20°33.4´w). small river south of zacken berg forskningsstation. the name is used by visiting scientists. kildeelven 75ø (75°15.0´n 20°57.4´w). name occasionally used by danish hunters for the river draining kildedal, also occasionally seen in the form lakseelven. kiledal 71ø-387 (71°23.3´n 27°38.5´w; map 4). wedge-shaped side valley to martin karlsen dal, th. sørensen land. named during the 1967–72 ggu scoresby sund expeditions (kile = wedge). kiledal 73ø-430 (73°15.8´n 25°28.8´w; map 4). wedge-shaped steep valley on western ymer ø, south of blomsterbugten. named during lauge koch’s 1947–49 expeditions by silvio eha. kilen 73ø-50i (73°59.0´n 21°24.1´w). minor wedge-shaped feature at the head of river 14 on the north slope of stensiö plateau, nw hold with hope. named by eigil nielsen during the 1931–34 tre års ekspeditionen. kilen 76ø-326 (76°44.0´n 24°39.0´w; map 4). wedge-shaped land area projecting north on the south side of borgjökel. named by the 1952–54 british north greenland expedition. kilen 81ø-72 (81°12´n 13°30´w; maps 1, 4). wedge of land on the east side of flade isblink, kronprins christian land. named by eigil nielsen during the 1938–39 mørkefjord expedition for its shape (kile = wedge). kilen fjelde 81ø (81°19.6´n 14°13.9´w). range of hills at the extreme nw of kilen, kronprins christian land. the name is found on a coloured geological map of kilen printed in 1991 (peder sen 1991). 229 kilesø 71ø-294 (71°58.1´n 26°41.4´w). lake in frederiksdal, nathorst land, dammed by a glacier and named by hans zweifel during lauge koch’s 1954–55 expeditions for its wedge-like shape. it is not present on recent aerial photographs. killingen 73ø (73°57.5´n 21°09.2´w). small island at the south end of stille ø in the finsch øer group. so named on an nsiu map (1932a), for its relative size (killingen = the kitten). kilmory fjeld 71ø-329 (71°43.7´n 25°11.9´w; map 5). mountain peak about 2100 m high between jupiter gletscher and pegasus gletscher, stauning alper. first climbed by john hunt’s 1960 expedition, and named kilmory, after the scottish base of the ‘national association of mixed clubs’ that had sponsored the expedition. kilroy 71ø (71°40.5´n 25°00.8´w; map 5). mountain peak about 1520 m high on the north side of mercurius gletscher, southern stauning alper. first climbed by james clarkson’s 1961 expedition. kilvrough fjeld 71ø-337 (71°44.3´n 24°57.3´w; map 5). mountain 2081 m high on the north side of bersærkerbræ, north stauning alper. first climbed by john hunt’s 1960 expedition, and named kilvrough, probably after ‘kilvrough manor outdoor education centre’. kilöya 73ø (73°58´n 21°10´w). name used on an nsiu map (1932a) for the present stille ø in the finsch øer group. named for its wedge-like shape (kil = wedge). kindtænderne 70ø-242 (70°54.7´n 21°49.5´w). row of summits on the east coast of liverpool land between randers fjord and horsens fjord, west of holloway bugt. named during the 1931–34 treårsekspeditionen by laurits bruhn for the resemblance to a row of teeth. king eider fjell 72ø (72°30.6´n 23°56.7´w). name used by the 1974 joint biological expedition for a hillside west of karupelv, sw traill ø. named after the king eider. kingofjeldet 74ø-337 (74°44.9´n 20°10.5´w; map 4). mountain on se kuhn ø. the name was proposed by the place name com mittee in 1939, and commemorates thomas kingo [1634–1703], a danish poet and bishop noted for his revised hymn book and more than 100 hymns. kings tinde 72ø-501 (72°05.0´n 25°16.3´w; map 5). peak 2470 m high on the sefström gletscher – gully gletscher divide, overlooking alpefjord. climbed by the 1963 cambridge expedition, and named after king’s college, cambridge (the king’s college of our lady and st. nicholas) founded in 1441. king’s college chapel, built 1446–1515 is regarded as the crowning glory of cambridge university. (kings, kings peak, picco king ). kingua i gaasefjord 70ø (70°06.0´n 28°00.0). name used for the inner part of gåsefjord in a report by hartz (1895) on the work of carl ryder’s 1891–92 expedition. it was probably intended as a descriptive rather than a formal place name, as ‘kingua’ in green landic signifies the inner part of a fjord. kirchenpauer bugt 74ø-74 (74°14.5´n 20°20.0´w). broad indentation of the ne coast of clavering ø on the south side of young sund. named by karl koldewey’s 1869–70 expedition as kirchen pauer bai, after gustav heinrich kirchenpauer [1808–87], businessman, politician and mayor of hamburg in 1870. he con tributed one of the zoology chapters to koldewey’s narrative. the bay is much less pronounced than shown on koldewey’s maps. norwegian hunters have used clavering bukta for the same feature. (kirchenpauers bugt, kirchenpauer bay.) kirkbrae 72ø (72°00.4´n 25°05.7´w; map 5). minor glacier on the ne side of sefström gletscher, stauning alper. kirkehytten – see domkirken. kirken 71ø-2 (71°07.0´n 21°53.6´w; map 4; fig. 53). mountain 1209 m high north of storefjord, liverpool land. named by william scoresby jr. in 1822 as church mount for its striking resemblance to a church. scoresby describes it as having two vertical towers at the summit with gable-formed tops, closely studded with pinnacles. the mountain was relocated in 1923 by henning bistrup during the voyage of the teddy, although he used the name biskop joseph fjeld. (church mountain, kirchberg, kirke bjerg, kirkefjellet.) kirkeruden 73ø-654 (73°35.0´n 24°37.8´w). feature in a cliff in south strindberg land, where a black rock with the shape of a church-window occurs in a light-coloured cliff. named by th. jo hansen during the 1931–34 treårsekspeditionen. kirkespiret [napassorssuaq] 74ø-40 (74°41.2´n 18°31.6´w; map 2). mountain 497 m high on lille pendulum. named by karl kolde wey’s 1869–70 expedition as kirchenspitze, because the rocky summit was reminiscent of a church spire. (church point.) kirriemuir 71ø (71°40.0´n 25°23.1´w; map 5). mountain 2100 m high at the head of jupiter gletscher, south stauning alper. first climbed by james clarkson’s 1961 expedition, and named after the small scottish town of the same name. kirschdalen 72ø-112 (72°33.8´n 24°53.2´w; map 4). valley in eastern lyell land draining east to kong oscar fjord. so named by eugène wegmann during the 1931–34 treårsekspeditionen, after swiss cherry brandy (= kirsch). kisbjerg 74ø-131 (74°16.4´n 20°51.5´w). mountain about 1369 m high on clavering ø. named by lauge koch’s 1929–30 expeditions as mt. kis, after a considerable outcrop of pyrite ore. (kisfjellet, kis bjerg.) kishmul borg 72ø-373 (72°04.2´n 24°39.5´w; map 5). mountain fig. 53. mirage view of the twin towers of kirken on liverpool land viewed from the east. the photograph was taken from a cruise ship and illustrates the results of a temperature inversion (dense, cold air beneath relatively warm air) that leads to distant objects towering above their normal height. kirken is only 1200 m high. spectacular arctic mirages with great vertical exaggeration are known as fata morgana. photo: c. kent brooks. 230 2450 m high at the head of kishmul gletscher, north stauning alper. named as kishmulborg by malcolm slesser’s 1958 expedition, probably after the legendary 14th century pirate who plied his trade on the ne coast of scotland. the mountain was first climbed by the 1963 imperial college expedition. kishmul gletscher 72ø-374 (72°05.8´n 24°28.4´w; map 4). glacier ne of kishmul borg, north stauning alper, that merges with skelbræ. named kishmul glacier by malcolm slesser’s 1958 expedition, although in an early report of the expedition it had been called glacier 21. kista dan gletscher 69ø-79 (69°57.0´n 27°36.0´w). smaller of two large glaciers draining into gåsefjord. named by w. stuart watt during the 1967–72 ggu scoresby sund expeditions. the kista dan (fig. 54) was the first of a series of ice-strengthened polar expedition and cargo vessels built by the j. lauritzen shipping company, and the 1100-ton vessel was initially used for the transport of lead ore from the mine near mesters vig. sailing under the name martin karlsen, it was the expedition ship of the 1968 ggu scoresby sund expedition. the same ship, renamed benja min bowring, was used as the support vessel for the 1979–82 transglobe expedition led by ranulph fiennes. see also martin karlsen bugt and magga dan gletscher. kista ø 72ø-329 (72°45.0´n 22°56.9´w; map 4). island in vega sund, between traill ø and geographical society ø. the name was proposed by søkortarkivet in 1956–57 following surveying of the channel through vega sund as an alternative approach for ships en route to mestersvig, and was given for the kista dan (fig. 54). see kista dan gletscher. grindøya has also been used. kjelbotn 73ø (73°06.6´n 23°00.0´w). norwegian hunting station about 1 km north of kap humboldt on se ymer ø, built by arktisk næringsdrift in 1929. it was named after olav kjelbotn [1898– 1966], a noted norwegian cross-country skier, who hunted in the region from 1929 to 1931 and built the station with ingwald strøm. kjelbotn made a memorable 70 km ski journey from kap hum boldt to myggbukta in deep snow in 32 hours. the station was intermittently manned in the periods 1929–41 and 1947–53, and has commonly been referred to as kap humboldt or humboldt. subsequently kjelbotn took part in the 1932–33 riiser-larsen antarctic expedition. (kjellbotn.) kjellbotn 72ø (72°55.3´n 23°47.7´w). small valley on west geo graphical society ø draining south into vega sund. so named (incorrectly with ‘ll’) on the nsiu maps of lacmann (1937) after olav kjelbotn. see also above. kjeldstrups tinde 71ø (71°53.2´n 25°08.9´w). summit about 2250 m high on the north side of roslin gletscher, between fimbulbreen and valhallbreen. it was climbed and so named by the 1996 norwe gian stauning alper expedition after øystein kjeldstrup [1956– 1976], a promising climber who died in a mountaineering accident. (kjeldstrups topp.) kjerulf fjord 72ø-417a 73ø-509 (73°03.0´n 27°22.4´w; map 4; see also fig. 65). n–s-trending fjord on the south side of innermost kejser franz joseph fjord. named during karl koldewey’s 1869–70 expedition, although the name is only found on the maps in payer’s (1876) narrative. probably named after theodor kjerulf [1825–88], professor of geology at the university in christiania (now oslo), and founder in 1858 of the geological survey of norway. a.g. nathorst observed in 1899 that payer’s kjerulf fjord did not exist in the position indicated and transferred the name to the present fjord farther west. josef hammar reached the inner end of the fjord by canoe in august 1899. the north half of the fjord is filled by stranded icebergs derived from nordenskiöld gletscher; louise boyd counted 525–530 large bergs here on a visit in 1931. (kjærulffjorden, kjer ulfs fjord.) kjerulfsdalen 72ø (72°53.8´n 27°33.4´w). name used by boyd (1932) in her report on her 1931 expedition for the present bocks rietdalen, south of the head of kjerulf fjord. kjovedammen 74ø (74°28.0´n 20°35.7´w). minor locality sw of zackenberg forskningsstation. the name has been used by visiting scientists (kjove = skua). kjoveland 71ø-353 (71°20.0´n 24°51.5´w). land area on the north side of the mouth of nordvestfjord, bordered to the east by schuchert dal. named by the 1963 geodætisk institut expedition after the long-tailed skua (= kjove). kjovestenen 74ø (74°28.6´n 20°35.7´w). minor locality north of zackenberg forskningsstation. the name has been used by visiting scientists. kjoveøen 74ø (7428.7´n 20°35.8´w). locality near zackenberg forsk ningsstation. the name has been used by visiting scientists. kjærsundet 73ø (73°05.2´n 23°02.0´w). sound between the se cape of ymer ø and robertson ø. used on the nsiu maps of lacmann (1937), the name was given for rolf kjær [b. 1897], a norwegian hydrographer who participated in nsiu expeditions to svalbard and east greenland, and from 1936 to 1967 was director of norges sjøkartverk. klassischer pingo 71ø (71°47.6´n 23°49.2´w). informal name used by müller (1959) in his report on work during lauge koch’s 1954–55 expeditions, for a pingo in pingo dal, northern jameson land. the pingo is of classic shape, 20 m high and with a circumference of 410 m. klatten 72ø (72°49.3´n 22°54.7´w). island in vega sund north of gåseøen. so named on the nsiu maps of lacmann (1937) for the shape (klat = lump). kleine kederbacher spids 71ø (71°52.9´n 25°36.3´w). mountain about 2400 m high on the west side of spærregletscher. named and first climbed by the 1967 berchtesgadener expedition. kleine sirius-pass 71ø (71°57.4´n 24°03.4´w; map 5). broad col at the head of the north branch of sirius gletscher between bellevue and taget, werner bjerge. the name is used in a description of climbing activities during lauge koch’s 1950 expedition (styger 1951). kleine sydney gletscher 71ø (71°56.7´n 25°37.7´w). name used by the 1967 berchtesgadener expedition for a tributary glacier on the west side of spærregletscher, stauning alper, which is more usually known as pollux glacier. named after sydney tinde at the head of the glacier. klinten 70ø-272 (70°06.0´n 23°17.9´w). cliffs on volquaart boon kyst between milano gletscher and østre borggletscher. so named during the 1931–34 treårsekspeditionen by laurits bruhn (klinten = the cliff ). klippedal 73ø-307 (73°47.6´n 23°10.0´w). valley in central hudson land draining into ankerbjergselv. the name was adapted from a suggestion by heinrich bütler during lauge koch’s 1936–38 expeditions (klippe = rock, cliff ). klipperne 74ø-323 (74°00.0´n 22°55.5´w). mountain range in north hudson land. named by heinrich bütler during lauge koch’s 1936–38 expeditions. klippeø 74ø (74°20.1´n 20°22.9´w). name used for basaltø in young sund in the ornithology report of løppenthin (1932). (klippeøen.) klitdal [kangerterajittap ilinnera] 70ø-118 (70°59.4´n 22°29.0´w; maps 3, 4). valley between liverpool land and jame son land, named by g.c. amdrup’s 1898–1900 expedition as klitdalen for the sand dunes (= klitter) in the southern part of the valley. (klit valley.) kloksethytten – see slippenhytten. kloksetøyane 72ø (72°43.2´n 22°47.6´w). small islands in vega sund, nw of silja ø. used only on nsiu maps (lacmann 1937), and named after ole klokset [b. 1910], a norwegian hunter who led an expedition to east greenland from 1933 to 1935. klosterbjerg 73ø (c. 73°18´n 29°07´w). name used for a mountain in the martin knudsen nunatakker during lauge koch’s 1951 expedition (buess 1953). like nearby spalenbjerg, it was named after a locality in the old town centre of basel, switzerland. 231 klosterbjerge 72ø-312 (72°14.5´n 25°57.3´w; map 5). mountain massif on the sw side of schaffhauserdalen. named by john haller following explorations during lauge koch’s 1954 expedition, after part of the old town centre of basel. klubben 70ø (c. 70°26´n 26°45´w). the name has been used for a mountain on eastern gåseland, west of falkepynt. klubben 74ø (c. 74°16´n 19°23´w). name used by the 1908–09 floren expedition for a feature in the vicinity of kap borlase warren (brandal 1930). exact position uncertain. klubtinde 71ø (71°47.3´n 25°24.1´w; map 5). mountain about 2550 m high on the ne side of orion gletscher. climbed by the 1996 norwegian stauning alper expedition, and named after the norsk tindeklub (a norwegian mountaineering club). klumpen 70ø (70°31.7´n 28°36.3´w). mountain between rolige bræ and vestfjord, the present rundefjeld, so named in helge vedel’s diary of carl ryder’s 1891–92 expedition (gulløv 1991). klus 73ø-313 (73°49.9´n 22°58.7´w). pass in central hudson land at the west end of dybendal. named by heinrich bütler during lauge koch’s 1936–38 expeditions. the name signifies a narrow valley or pass, and is commonly used for the narrow valleys in the limestone country of the jura, switzerland. klægbugt 77ø-70 (77°36.5´n 20°47.3´w). bay on the east coast of nordmarken, innermost skærfjorden. so named by david malm quist during the 1931–34 treårsekspeditionen, because the coastal flats are of clay which when wet is so sticky that progress is impossible (klæg = sticky). kløft 1 74ø (74°25.1´n 20°14.9´w). small ravine, the northern upper branch of sandstensdal, western wollaston forland. used as a re ference locality by rosenkrantz (1932). kløft i 76ø-263 (76°22.7´n 18°41.9´w; map 4). narrow ravine on the east side of store koldewey. named by the 1906–08 danmarkekspeditionen, and first used as a geological reference locality by ravn (1911). håkon jarner used vardekløft for the same feature in june 1907 (j. løve, personal communication 2009). kløft 2 74ø (74°24.8´n 20°15.2´w). small ravine, the southern upper branch of sandstensdal, western wollaston forland. used as a re ference locality by rosenkrantz (1932). kløft ii 76ø-264 (76°22.3´n 18°41.6´w; map 4). narrow ravine on the east side of store koldewey, a little south of kløft i. named by the 1906–08 danmark-ekspeditionen, and first used by ravn (1911) as a geological reference locality. kløftbjerge 71ø-352 (71°20.0´n 25°40.0´w). mountain range with a summit ice cap in ne renland, south of the mouth of nordvest fjord, noted for its many ravines (kløft = ravine). named by the 1963 geodætisk institut expedition. kløftdalen 73ø-701 (73°12.0´w 27°04.0´n). narrow, cleft-like valley in frænkel land. named by john haller following explorations during lauge koch’s 1949–51 expeditions. kløftelv 70ø-124 (70°54.0´n 22°37.5´w). river nw of the head of hurry inlet. named during lauge koch’s 1926–27 expeditions by alfred rosenkrantz, originally as corrie river, because it drains a glacial feature, a valley formerly occupied by a glacier and known as a corrie. the name kløft (= ravine) is an alternative rather than a translation of corrie. kløftelv 72ø-128 (72°52.4´n 25°05.6´w). river on nw ella ø, draining from ulvesø into solitærbugt. so named by the ella ø wintering party 1931–32, during the 1931–34 treårsekspedi tionen, because it drains through a ravine. kløftelv 73ø-361 (73°48.2´n 24°53.4´w). river in central strind berg land that drains from søgletscher through a ravine into laksesø in brogetdal. named during lauge koch’s 1948–49 expeditions by hans r. katz. kløftfjeld 76ø-250 (76°51.9´n 19°29.7´w). hillside on winge kyst between snenæs and lille snenæs, cut by a ravine. named by the 1906–08 danmark-ekspeditionen in the form kløftfjeldet, and first used by lundager (1912) in his description of the vegetation of the region. kløftgletscher 74ø-322 (74°37.9´n 22°14.5´w). glacier on the sw side of tyrolerdal, named by louise boyd’s 1937 expedition as kløft glacier because it occupies a steep and narrow ravine. on some maps the names of copeland gletscher and kløftgletscher have been interchanged. (kløft gletscher.) kløfthytten 77ø (77°15.4´n 19°25.4´w). hut built in 1940 by the mørkefjord expedition on the west side of fladebugt, north of fig. 54. the ice-strengthened cargo vessel kista dan in the 1950s, on its way to nyhavn near mesters vig to pick up lead-zinc ore. the john haller photograph collection, geus archive. 232 michelangelo kløft, germania land. it is also known as knuths minde and kap li hytten. kløvskæret 76ø-272 (76°46.0´n 18°24.0´w). small island on the east coast of germania land, north of øksebladet. named by the 1938–39 mørkefjord expedition (kløve = cleave, split). knabendalen 72ø (72°50.8´n 22°40.0´w). valley on south geo graphical society ø between lysdal and adam af bremen dal. the name is used on the nsiu maps of lacmann (1937), and was given for the norwegian zoologist nils knaben [1898–1969], who participated in the 1929 and 1930 nsiu expeditions. he became head curator at the zoological museum in oslo. knacke glacier 72ø (72°03.3´n 25°12.1´w; map 5). minor glacier on the ne side of sefström gletscher. knasten 73ø-196 (73°42.4´n 21°33.2´w; map 4). mountain 768 m high on the east side of loch fyne, western hold with hope. it has also been called øienfjellet. knasten 76ø-216 (76°41.5´n 21°58.4´w). southernmost solitary island in borgfjorden. named during the 1938–39 mørkefjord expedition by paul gelting in november 1938, for its appearance as a spot or pimple on the otherwise level fjord ice. knebel vig 72ø-84 (72°16.2´n 22°18.4´w; map 4). bay on the south side of mountnorris fjord, se traill ø. named during the 1931–34 treårsekspeditionen by ove simonsen after the danish locality of the same name in the mols district of jylland, denmark. kneet 73ø (73°33.5´n 22°46.0´w). name occasionally used by norwegian hunters for the pronounced bend towards the eastern end of moskusoksefjord (= the knee). knibtangen 73ø-372 (73°47.7´n 26°15.5´w). two glacier tongues in eremitdal in northern andrée land, that descend from opposite sides of the valley and almost meet, and have a pincer-like shape in plan (knibtangen = the pincers). named during lauge koch’s 1948–50 expeditions by erdhardt fränkl. knighton fjord 69ø-10 (69°21.0´n 24°38.0´w; map 3). fjord on the northern blosseville kyst. william scoresby jr. in 1822 gave it the name knighton bay in honour of sir william knighton [1776–1836], physician to the prince of wales, who in 1822 was appointed private secretary and keeper of the privy purse to george iv. ejnar mikkelsen had suggested c. holms bugt for the same feature in 1924. (knighton bugt, knighton bai.) knivbjerg 73ø-689 (73°26.8´n 26°33.8´w; map 4). mountain in sw andrée land. named by john haller, following explorations during lauge koch’s 1949–51 expeditions, for its sharp snow ridge (kniv = knife). knivneset 74ø (c. 74°16´n 19°23´w). name used by the 1908–09 floren expedition (brandal 1930) for a peninsula or ridge in the vicinity of kap borlase warren (kniv = knife). exact position uncertain. (knivberget.) knivodden 72ø (72°01.1´n 23°03.7´w). peninsula on the se side of the mouth of antarctic havn. the name is found on norsk søkort 511 (1937) and in den grønlandske lods (1968). knoen 73ø-549 (73°00.7´n 27°58.5´w). mountain 2300 m high in northern goodenough land, named by j.m. wordie’s 1929 expedition as knuckle. knogledal 71ø-412 (71°58.2´n 23°05.5´w). side valley to flexur dal, south of antarctic havn. named by katharina perch-nielsen during the 1967–72 ggu scoresby sund expeditions for the abundance of bones (= knogler) of musk ox. knolden 72ø-180 (72°55.1´n 22°51.6´w). one of the peaks of julekagen, geographical society ø. the name was one of a group of names given by the place name committee in 1939 and was given for the shape (knold = hill). meydenbauerfjell has also been used. knolden 73ø-410 (73°53.1´n 29°03.1´w). nunatak in j.l. mowinckel land. the name appears on a map in katz (1952) as knalden, and was corrected to ‘knolden’ by the place name com mittee. knolden 73ø-443 (73°15.1´n 22°24.0´w). minor summit on the south flank of knuden, near knudedal, se gauss halvø. named during lauge koch’s 1950 expedition by p. graeter. knolden 73ø-554 (73°01.2´n 27°52.4´w). mountain 2302 m high ne of mercanton gletscher in north goodenough land, named by j.m. wordie’s 1929 expedition as the knoll. knolden 74ø-227 (74°01.7´n 21°37.0´w). minor feature in nw hold with hope, between river 6 and river 7, on the north slope of frebold bjerg. named by eigil nielsen during the 1931–34 treårs eks peditionen. blokken has been used for the same feature. knolden 74ø-307 (74°05.7´n 21°16.5´w). small promontory im mediately below eskimonæs station, which interrupts the line of the beach in østhavn. the name originated from the wintering party at eskimonæs during the 1931–34 treårsekspeditionen (knold = hill). knolden 74ø-347 (74°24.5´n 19°18.5´w). hill of basalt in lower dronning augustadalen, wollaston forland. named by wolf maync and andreas vischer during lauge koch’s 1936–38 expeditions. knolden 76ø (76°45.7´n 18°48.5´w). name used by the 1906–08 danmark-ekspeditionen for a minor feature on the south coast of germania land nw of bådskæret. the name is found on a handcoloured map of the danmark havn area in the arktisk institut, copenhagen (j. løve, personal communication 2009). knophdalen 72ø (72°55.3´n 24°18.5´w). valley on west geo graphical society ø draining south into vega sund. so named on the nsiu maps of lacmann (1937) after gunnar knoph [b. 1898], a norwegian hunter. see also knophstua. knophstua 73ø (73°55.0´n 20°54.6´w). norwegian hunting hut on the north coast of home forland south of terneskær, about 1 km se of the mouth of rødelv, built by finn devold’s expedition in 1928. it replaced a hut built by the foldvik expedition in 1926, but taken down in 1927. this name appears on an nsiu map (1932a). the hut was named after gunnar knoph [b. 1898], who built the hut and hunted in the region from a main base at örnereiret from 1929 to 1930. it has also been known as rødelv. (knoph-stua, knopstua.) knorten 73ø (73°42.4´n 20°34.5´w). hill 292 m high on the east coast of hold with hope. so named on the 1932a nsiu map because of its shape, a knobby lump. knot hill 72ø (73°32.5´n 23°58.6´w). name used by the 1974 joint biological expedition for a hill west of karupelv, sw traill ø. it was named after the knot, a small wader. knotten 72ø (72°51.7´n 21°45.9´w). small island off the coast of east geographical society ø, sw of kap mackenzie. used only on nsiu maps (lacmann 1937), the name was given for the shape (knott = rounded lump, protuberance). knud rasmussen land 68ø, 69ø, 70ø (68°20´n–70°n). extensive land area between the south coast of scoresby sund and kanger lussuaq. the name appears throughout the official report of the 7th thule expedition (gabel-jørgensen 1940), and was the officially approved name for this region from 1940 until 1953. knud johan victor rasmussen [1879–1933], a noted danish–green landic explorer and ethnographer, died shortly after the return of this expedition, in the course of which much of the region was photographed from the air. in 1953 the name was transferred to cover north greenland between melville bugt and danmark fjord, an area including thule (now dundas), the base from which knud rasmussen organised many of the thule expeditions. however, the name knud rasmussen land is still often used in its original sense for the region south of scoresby sund, especially by mountaineering expeditions. knud ringnes nunatak 73ø-574 (73°44.8´n 29°41.9´w; map 4). nunatak north of evers gletscher. named by arne høygaard and martin mehren in 1931 after knud ringnes, a norwegian businessman and director. the ringnes brewery was at one time the 233 largest in norway, and was noted for its support for norwegian arctic exploration. as chairman of the fram committee, knud ringnes was responsible for the preservation of the fram as a museum ship in 1936. knudedal 73ø-440 (73°16.0´n 22°22.1´w). steep valley draining south from knolden, se gauss halvø. named during lauge koch’s 1950 expedition by p. graeter. (knuden dal.) knuden 73ø-100 (73°17.1´n 22°18.4´w). mountain north of kap franklin, se gauss halvø. named during the 1931–34 treårs ekspeditionen by th. johansen (knude = knot). franklinfjellet has also been used. knuds dal 75ø (75°08.8´n 19°52.6´w). name occasionally used by danish hunters for a minor valley in sw hochstetter forland east of niels hansen næs (nyholm-poulsen 1985). personal name. knudsens nunatakker – see martin knudsen nunatakker. knudshoved 73ø-72 (73°43.9´n 20°27.1´w; maps 2, 4). peninsula on the east coast of hold with hope. the name is credited to the ship’s crew aboard the godthaab in 1930, and was given for its supposed similarity to the peninsula of the same name near ny borg, denmark. (knuts hode.) knudshoved station 73ø-73 (73°42.5´n 20°32.2´w). danish hunting station on the east coast of hold with hope, 3 km south of the peninsula knudshoved. it was manned in the periods 1930–32, 1934–40 and 1945–46. the station has usually been referred to simply as knudshoved. it was built by nanok in 1930 as a replacement for the carlshavn station, burnt down in 1927. knuthsminde 77ø (77°15.4´n 19°25.4´w). hut built in 1940 for the mørkefjord expedition on a peninsula on the west side of flade bugt, and named after eigil knuth, leader of the expedition. the hut has also been known as kløfthytten and kap li hytten. knytlingen 73ø (73°38.3´n 20°27.2´w). small skerry very close to the coast of eastern hold with hope. so named on an nsiu map (1932a), the name derives from the norwegian word for something nearby or connected, in this case presumably the closeness of the skerry to the coast. knækdalen 73ø-606 (73°12.9´n 27°55.4´w; map 4; see also fig. 65). valley in sw frænkel land. so named by ove simonsen during the 1931–34 treårsekspeditionen, because of the right-angled bend (= knæk) in the valley. gregory valley has also been used. (knægt dalen, knækdal, knäkdalen.) knækelven 73ø-622 (73°11.6´n 27°39.8´w). river flowing in knækdal, sw frænkel land. the name was proposed by the place name committee in 1935. gregory river has also been used. knækelvhytten – see bræhytten. knækket 70ø-391 (70°16.0´n 26°42.0´w; maps 3, 4). relatively narrow part of gåsefjord where the fjord bends and changes direction (knækket = the break). named by the 1963 geodætisk institut expedition. knæksø 75ø-108 (75°53.3´n 22°00.0´w; map 4). lake in western nørlund land, draining the north branch of ejnar mikkelsen glet scher. the name is a modification of an original suggestion by lauge koch, approved in 1957, and records the pronounced bend in the lake. knøsen 77ø-93 (77°48.2n 19°27.6´w; map 4). mountain on se gamma ø with a broad summit ice cap. named by the 1938–39 mørkefjord expedition after a danish locality of the same name. kobberpynt 70ø-11 (70°31.0´n 28°21.0´w). peninsula on the north coast of vestfjord. named in this form by the carl ryder 1891–92 expedition because the deep, red-brown weathering, ultrabasic in tru sions appeared to contain copper ore – there are small amounts of titanomagnetite and pyrrhotite. this is probably the same locality as nordenskjöld’s (1907) black point or sorte pynt. see also sortepynt. the variation kalebarodden that appears in helge vedel’s published diaries of carl ryder’s 1891–92 expedition (gulløv 1991) is a transcription error for kobberpynt (j. løve, per sonal communication 2010). koch stones 76ø (c. 76°42.0´n 18°33.0´w). this designation is used in some of the the 1906–08 danmark-ekspeditionen reports for the cairn at kap bismarck used as a surveying station. the green landers on the expedition used the term to distinguish cairns built by members of the expedition (e.g. j.p. koch) from the various constructions built by the inuit that were known as ‘eskimo stones’. koch’s corridor 77ø (c. 77°25´n 23°00´w). term used by the 1952–54 british north greenland expedition for the route across storstrømmen from near annekssøen to ymer nunatak, a problem-free passage used by j.p. koch during the 1906–08 danmarkekspeditionen. the british expedition failed to find it in 1953 while exploring for a route for their weasel tractors, probably be cause this glacier periodically surges, with resulting major changes in surface features. kocheler spids 71ø (71°50.3´n 25°17.0´w; map 5). mountain on the sw side of roslin gletscher. climbed by karl herligkoffer’s 1966 expedition on 21 august, and named after kochel, a small town on the banks of the kochelsee in the bavarian alps. kochi-ridge – see western upper terrace. kochsvighytten 75ø-102 (75°15.9´n 19°58.9´w). danish hunting hut in lauge koch vig on the sw coast of hochstetter forland, built by nanok in september 1931. (koch vig hytten.) kodak ridge 70ø (70°56.3´n 25°52.1´w). name given to a 1500 m high ridge in northern milne land by the 1989 greenland milne land expedition. kofoed-hansen bræ 77ø-37 (77°32.0´n 21°44.0´w). large glacier flowing north between sønderland and nordmarken. named by the 1906–08 danmark-ekspeditionen as kofoed-hansens bræ, after otto joachim moltke kofoed-hansen [1854–1918], a director of the danish admiralty who had shown great interest in the expedition. kohleninsel 74ø, 75ø (74°50.4´n 20°15.3´w). temporary name given during karl koldewey’s 1869–70 expedition to the present kuhn ø, and so named because layers of coal up to 50 cm thick were found at kap hamburg. kokkens lyst 71ø (71°36.4´n 22°36.3´w). hut in the inner part of nathorst fjord erected in the summer of 1977 by jan juel-brock dorff for mestersvig airfield. it was intended as an emergency hut for aircraft personnel operating between scoresbysund and mesters vig. brockdorff was cook (= kok) at mestersvig airfield. kolddal 72ø-520 (72°23.0´n 23°00.0´w). small valley on se traill ø draining into fossdal. named by geoffrey halliday following botanical work during the 1961 leicester and 1971 northern universities expeditions. koldewey øer 76ø-38a (76°28.0´n 18°50.0´w). island group in cluding store koldewey and lille koldewey. named by karl kolde wey’s 1869–70 expedition as koldewey inseln after karl kolde wey [1837–1908] (fig. 55), captain of the germania and leader of the first and second german polar expeditions. the present store koldewey is shown as three large islands on kolde wey’s maps. there is a remote possibility (tornøe 1944) that the island ‘ran sey’ on a 1706 map by torfæus might correspond to this island group. (koldewey island, iles koldewey, koldewey islands.) kolding fjord 70ø-224 (70°42.8´n 21°39.0´w; map 4). fjord on the east coast of south liverpool land, so named during the 1931–34 treårsekspeditionen by laurits bruhn after the fjord of the same name in jylland, denmark. kolledalen 71ø-150 (72°00.0´n 23°21.0´w). valley west of antarctic havn in north scoresby land. derived from the kolldal of norwegian hunting expeditions based at karlsbak, antarctic havn. the name is found in the form kolldalen on norsk søkort 511 (1937). antarctic dal was used by some members of lauge koch’s expeditions. kollen 70ø-127 (70°51.1´n 22°47.2´w). mountain on the west side of the head of hurry inlet between postkassen and statuebjerg, and joined to statuebjerg by a high col. named by alfred rosen 234 krantz and tom harris during lauge koch’s 1926–27 expeditions as coll mountain. kollen 71ø-90 (71°36.2´n 22°22.1´w). elongate hill on canning land, so named during the 1931–34 treårsekspeditionen by arne noe-nygaard, because of its shape (kolle = a rounded top). koloni 70ø (70°29.1´n 21°57.9´w). name occasionally used on maps and in publications for the town of illoqqortormiut / scoresbysund (koloni = colony). kolossen 72ø-301 (72°01.3´n 24°02.3´w). mountain 1038 m high between mellem gletscher and østre gletscher, northern werner bjerge. named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk, and climbed in 1953 by wenk. it ap peared on earlier maps of styger (1951) as centralen. kolstad 75ø (75°37.0´n 19°30.1´w). norwegian hunting station south of haystack on the east coast of hochstetter forland. it was built by john giæver’s expedition in 1932, and has also been known as ottostrand. the name kolstad has also been used for another norwegian station farther north, also erected in 1932 – see jónsbú. the name kolstad was intended as a tribute to the norwegian prime minister, but he considered the political implications unfortunate, and alternative names were subsequently used. kolthoffhytten – see johnsenhytten. kommafjæld 73ø (73°29.8´n 20°32.0´w). name suggested by gustav thostrup for midterfjeld in sw hold with hope. during the voyage of the teddy in 1922 thostrup sketched the coast and observed that a glacier on the east side of the mountain had the shape of a comma. kong christian ix land 69ø-70, 70ø-89 (65°30´n–70°n). exten sive region between ammassalik and the south side of scoresby sund. the name had been given by gustav holm to the area north of ammassalik, and was later extended by g.c. amdrup following his 1898–1900 expedition as far as scoresby sund. the region between kangerlussuaq (c. 68°30'n) and scoresby sund has sometimes been referred to as nordlige christian den ix’s land (stor gaard 1927). it had become a tradition to name newly explored areas of greenland after the reigning monarch, and christian ix [1818–1906] was king of denmark from 1863. kong christian x land 70ø-90, 71ø-323, 72ø-485, 73ø-720, 74ø-405, 75ø-109, 76ø-346 (70°n–76°n). major geographical division of east greenland, with a south boundary following the south coast of scoresby sund, and a north boundary at bessel fjord. the region was overflown by lauge koch in 1932, and during the planning session for the flight koch is reported to have said “let’s get this over with as quickly as possible and then we can call the whole thing king christian x land.” the name was first used on the 1932 1:1 million scale geodætisk institut map, and follows the tradition of naming newly explored land areas for the ruling monarch at the time of exploration. christian x [1870– 1947] was king of denmark from 1912. kong frederik viii land 75ø-110, 76ø-244, 77ø-140a, 78ø-42, 79ø-25, 80ø-109, 81ø-127 (76°n–81°n). major geographical division of northern east and eastern north greenland, with a south boundary running along bessel fjord, and a north boundary through the middle of independence fjord and academy glet scher. the name was used on the 1906–08 danmark-ekspedi tionen maps for the region 79°–81°30´n, on a map by storgaard (1927), and came into general usage following the 1931–34 tre årsekspeditionen explorations. the region was explored largely by the 1906–08 danmark-ekspeditionen, the 1909–12 alabama ex pedition and j.p. koch’s 1912–13 expedition, when the ruling monarch was frederik viii [1843–1912], king of denmark from 1906. storgaard (1927) proposed a division of this extensive region into two parts (nordlige and sydlige frederik den viii’s land) with a division along nioghalvfjerdsfjorden. kong oscar arkipelag 72ø, 73ø (72°–74°n). this was one of the physiographic divisions of east greenland proposed by storgaard (1927), and was intended to cover the land areas and islands be tween latitudes 72° and 74°n. it approximately corresponds to the arctic riviera of hofer (1957). kong oscar fjord 72ø-54 (72°22.0´n 24°00.0´w; maps 3–5; see also fig. 78). major fjord 10–25 km in width, bounded by traill ø and geographical society ø to the east, and ella ø, lyell land and the stauning alper to the west. named by a.g. nathorst’s 1899 expedition as konung oscars fjord after oscar ii [1829–1907], king of sweden from 1872 to 1907, and a supporter of the expedition. white (1927) had suggested the name be restricted to only the n–s-trending section of the fjord, with a corresponding greater extent for davy sund, but this proposal has not been followed. (king oscar fjord, kong oscars fjord, fjord, fiord de roi oscar, kong oskarfjord.) kong wilhelm land 75ø-32 (c. 75°45´n 22°45´w; maps 2, 4; see also fig. 81). land area west of the head of bredefjord, between 75°25'n and 75°58°n. named by karl koldewey’s 1869–70 expedition as könig wilhelms land, after wilhelm i [1797–1888], king of prussia 1861–1888 and emperor of germany 1871–1888. he had made the largest single donation to the expedition finances, and the koldewey expedition reports (verein für die deutsche nord polarfahrt in bremen 1873–74) are dedicated to wilhelm i. kolde wey’s original usage was in a much broader sense than the present, covering an extensive unmapped region between latitudes 75°– 77°n, that appeared on the duke of orléans’ map from 1905 as terre du roi guillaume (fig. 9). (king wilhelm land, king wil liam’s land, kong vilhelms land, terre du roi guillaumo.) kongeborg 72ø (72°35.4´n 24°22.9´w). norwegian hunting hut built in 1932–33 for helge ingstad’s expedition on the sw point of traill ø, at the south end of the cliffs known as kongeborgen. the roof of the hut was an upturned boat. the hut was replaced in 1950 by a new hut, known as kongeborgen. (kongeborgenhytten, kongs borg.) fig. 55. karl koldewey [1837–1908], the leader of the 1869–70 second german north pole expedition to northern east greenland. from: verein für die deutsche nordpolarfahrt in bremen (1873–74). 235 kongeborgen 72ø (72°35.4´n 24°22.9´w). norwegian hunting hut built in august 1950 for hermann andresen’s expedition. it re placed the kongeborg hut on the same site. (kongeborghytte.) kongeborgen 72ø-55 (72°42.0´n 24°23.0´w; map 4; fig. 29). western cliffs of traill ø, which reach altitudes of 1300–1700 m. named konungaborgen by a.g. nathorst’s 1899 expedition for its impressive high walls and pyramid-formed tops and projections bordering kong oscar fjord. on his chart, nathorst (1900) used the form kungaborgen. (royal castle, konga borgen, king’s castle mountain.) kongespejlet 71ø (71°58.0´n 24°20.0´w ´w). glacier draining from the central stauning alper se and south to the head of schuchert dal, the present schuchert gletscher. the name was one of a group of names for glaciers given by the place name committee in 1939, which replaced proposals by hans stauber. the name was officially approved from 1939 to 1960, although it is only occasionally found on maps. in 1960 the name was replaced by the widely used schuchert gletscher. the kongespejlet is one of the icelandic manuscripts dating from c. 1250, in which greenland is described. konglomeratelv [kaporniagaqarpik] 71ø-226 (71°20.2´n 24°48.7´w). minor river draining south to the west side of nord ostbugt. named by enrico kempter during lauge koch’s 1956–58 expeditions, for the outcrops of conglomerate. konglomeratelv 74ø (74°38.3´n 20°41.7´w). river draining north into lindeman fjord, northern wollaston forland. the name was used by wolf maync during lauge koch’s 1936–38 expeditions, and given for the presence of conglomerate (maync 1947). konglomeratnæs 73ø-436 (73°02.2´n 24°40.9´w). peninsula on the south coast of ymer ø, between karl jakobsen bugt and bo tanikerbugt. it was named by silvio eha for the conglomeratic rocks (eha 1953). konglomeratpas 71ø (71°29.2´n 24°56.1´w). minor pass between gurreholm dal and konglomeratelv, on the west side of schuchert flod. the name was used by kempter (1961). konglomeratrücken 74ø (74°51.7´n 20°30.9´w). ridge on west kuhn ø west of baselbjerget, where maync (1947) reported finds of conglomerates during lauge koch’s 1936–38 expeditions. kongsholmen 77ø-68 (77°28´n 20°09´w). island in innermost skærfjorden, so named during the 1931–34 treårsekspeditionen by david malmquist after kungsholmen, an island in central stockholm, sweden. konrad bjerg 77ø-71 (77°38.2´n 20°41.1´w). mountain on the east coast of nordmarken, innermost skærfjorden. named by david malmquist during the 1931–34 treårsekspeditionen in the form konradsberg, after a swedish mental hospital of that name. ‘you are now ready to go to konradsberg’ was a rather usual comment to anyone who made a stupid remark. kontaktravine 74ø-166 (74°22.7´n 20°35.7´w). ravine on ne clavering ø. so named by arne noe-nygaard and gunnar sävesöderbergh during the 1931–34 treårsekspeditionen, because a geological boundary occurs here. koorajik 70ø-292 (70°27.7´n 22°16.8´w). stream east of ittaajim mit [kap hope], sw liverpool land. recorded by the 1955 geodætisk institut name registration, the name means ‘the little river’. (kôrajik.) koppefjellet 72ø (72°56.6´n 24°20.5´w). mountain 1730 m high on west geographical society ø, east of svedenborg bjerg. used on the nsiu maps of lacmann (1937), and named after karl koppe [1844–1910], german professor of geodesy at the technischen hochschule, braunschweig, who contributed important developments in photogrammetry. kopperneshuset 74ø (74°29.6´n 18°59.9´w). norwegian hunting station west of kap wynn, wollaston forland, built by the floren expedition in 1908, and named after an ålesund merchant, h. koppernes, who had helped finance the seven-man expedition. only the foundations of this house remain. the last timbers were used to build the hird expedition hut 300 m to the east in 1928. a new hut known as gåsneshuset was built beside it by arktisk næringsdrift in 1929. (koppernes-tufta.) kôrajik – see koorajik. koralbjerg 74ø-364 (74°21.0´n 20°47.6´w). mountain 1370 m high on ne clavering ø. named by wolf maync and andreas vischer during lauge koch’s 1936–38 expeditions for the finds of large numbers of fossil corals in permian dolomites. bundermannfjellet has also been used. koralkløft 73ø-89 (73°19.1´n 22°42.8´w). small ravine west of margrethedal on se gauss halvø. originally the river was named by lauge koch’s 1929–30 expeditions in the form coral creek or coral river, because of the abundance of fossil corals. sävesöderbergh’s (1934) maps suggest the official placing west of camp creek may be incorrect. korpvatnet 72ø (72°45.4´n 22°26.1´w). small lake sw of freycinet bjerg on sw geographical society ø. so named on the nsiu maps of lacmann (1937) (korp = raven). korridoren 70ø-409 (70°48.0´n 26°12.0´w; map 4; fig. 56). deep valley occupied by a major glacier extending from central øfjord eastwards across milne land. so named during the 1967–72 ggu scoresby sund expeditions by niels henriksen, because the valley provided a route, or corridor, across milne land that was often used by helicopters. korridoren 76ø-351 (76°44.6´n 18°48.4´w). broad sound between kap bornholm, the north point of lille koldewey and bådskæret, danmarkshavn, marked by a strong current. the name was reported by hans meltofte as in general use by weather station staff in 1969–71. korsbjerg 70ø-232 (70°42.5´n 22°00.0´w; map 4). mountain in southern liverpool land, so named during the 1931–34 treårs ekspeditionen by laurits bruhn, probably for the four summit ridges in the shape of a cross. a cairn at the 1400 m high summit was discovered by the 1989 snow dance expedition, and contained records of an ascent by two swiss members of lauge koch’s expeditions in 1933, and a visit by the geodetic institute in 1969. korsbjerg 72ø-206 (72°10.8´n 23°56.1´w; map 5). mountain in north scoresby land west of mesters vig, with the highest summit at 1060 m. so named by prospecting teams associated with lauge koch’s 1948–49 expeditions, because it is formed by several intersecting high ridges. korseneset 74ø (c. 74°16´n 19°23´w). name used by the 1908–09 floren expedition for a peninsula in the vicinity of kap borlase warren (brandal 1930). exact position uncertain. korsgletscher 74ø-390 (74°18.5´n 25°13.5´w; map 4). glacier in bartholin land dividing into three branches in a cross-shaped pattern. named by john haller following explorations during lauge koch’s 1956–58 expeditions. korsspids 72ø-321 (72°03.6´n 25°07.1´w; map 5; fig. 27). massive mountain about 2780 m high east of the head of cavendish gletscher, stauning alper. climbed by the 1963 university of cambridge expedition. hans gsellman’s 1957 expedition may also have climbed the peak, and called it weisswand, but no sign was seen on the summit, and it is suspected they climbed the adjacent peak to the se. so named during lauge koch’s 1954 expedition by john haller, probably because the ridges form a cross. korstrollet – see søstjernen. kortedal 71ø-316 (71°42.4´n 24°35.6´w; map 5). valley south of roslin gletscher draining east into schuchert dal. named by the place name committee in 1959 as a replacement for a suggestion by enrico kempter. kosmoceras bjerg 71ø (71°27.3´n 23°49.4´w). name found only in surlyk et al. (1973), where it was used for a mountain in jameson land west of olympen. it derives from finds of fossil kosmoceras ammonites. kosmocerasdal 70ø (70°44.5´n 25°29.1´w). minor valley on se 236 milne land draining ne into charcot bugt. the name appears on maps of callomon & birkelund (1980), and derives from finds of kosmoceras ammonites. it has also been called chattonkløft. kostenbaderbjerg 70ø-42 (70°42.0´n 25°19.6´w). minor peak 460 m high nw of kap leslie, east milne land. named during the 1931–34 treårsekspeditionen as kostenbader berg or kosten baderberg by hermann aldinger. origin uncertain, but possibly given for a german geologist. kote 800 70ø (70°39.4´n 25°56.9´w). prospectors name for the 800 m high isolated hill west of bay fjelde, east milne land, where nordisk mineselskab investigated a placer deposit. shallow drilling has proved 5 million tons of ore with 1–3.8% zr and 0.5– 1.9% rare-earth elements (harpøth et al. 1986). krabbedalen 70ø-353 (70°06.4´n 22°14.3´w). small valley at the head of bopladsdalen, kap brewster, where well-preserved fossil crabs were collected by d. mackney and f.w. sherrell during lauge kochs 1951 geological expedition. the name was used by hassan (1953). krabbegletscher 72ø-268 (72°01.0´n 25°26.9´w; map 5; fig. 38). glacier draining into dammen at the head of alpefjord, notable for the two partly submerged moraine ridges resembling the claws of a crab. named during lauge koch’s 1954 expedition by john haller. kradshytten 76ø (76°45.7´n 18°48.5´w). hut built by danmarks havn weather station personnel in 1968, ne of bådskæret, near danmark havn. the name derives from an expression in use at the station in the 1960s, which meant ‘to go hunting’. kragenrede 71ø (71°08.3´n 26°29.1´w). summit 2037 m high on the corner between edward bailey gletscher and catalinadal, ren land. climbed and named by the 2007 west lancashire mountain eering group expedition. kranges fjeld 75ø (75°18.3´n 21°15.1´w). mountain on the south side of ardencaple fjord between femdalen and kildedal, corresponding to the present vesterport. the name was used by nanok hunters (hvidberg 1932), and was said to have been given for a copenhagen lawyer of that name. kratersee 72ø (72°33.1´n 23°37.3´w). name used by fritz müller during lauge koch’s 1954–55 expeditions for a lake 100 m across in his kraterseepingo, karupelv valley, traill ø (kratersee = crater lake). kraterseepingo 72ø (72°33.1´n 23°37.3´w). name used by müller (1959) in his report on work during lauge koch’s 1954–55 expeditions, for a 12 m high pingo in karupelv valley, traill ø. kratersee occupies the centre of the pingo. kratlien 74ø-305 (74°05.3´n 21°16.5´w). slope on the se side of the peninsula of eskimonæs. the name originated from the wintering party at eskimonæs during the 1931–34 treårsekspedi tionen, and refers to an area covered by scrub (= krat). kravebjerg 71ø-277 (71°54.7´n 23°42.8´w). mountain in the se werner bjerge, named by peter bearth and eduard wenk during lauge koch’s 1953–54 expeditions for a collar-like feature (krave = collar). it was climbed by wenk in 1953. krebs bjerg 77ø-47 (77°13.0´n 24°23.9´w; map 4). mountain in dronning louise land, so named by the 1909–12 alabama expedition. probably named after holger klingberg krebs [1872– 1953], a danish marine officer promoted to captain in 1909 (j. løve, personal communication 2009). it was climbed by members of the 1952–54 british north greenland expedition, and their surveying station on the summit was informally referred to as lurcher’s crag. (krebs nunatak.) krebsedal 70ø-49 (70°42.0´n 25°18.1´w). small valley on the east coast of milne land between charcot havn and kap leslie. named by hermann aldinger during the 1931–34 treårsekspedi tionen as krebstal or krebs-tal, for the fossil crabs. (crab valley.) kriemhildbreen 74ø (74°25.0´n 21°06.6´w). glacier on north clave ring ø. so named on the nsiu maps of lacmann (1937), after krimhild, wife of siegfried, burgundian princess of worms who killed the defenceless hagen in the german epic poem from c. 1200, the nibelungenlied. kristianshavn – see christianshavn. kristiern nielsen dal 71ø (71°47.4´n 23°49.1´w). valley draining east into ørsted dal, the present pingo dal. the name was one of a group of names given by the place name committee in 1939, which replaced proposals by hans stauber. the name was officially approved from 1939 to 1957, although only rarely used on maps (e.g. hübscher 1943). kristiern nielsen was a priest who accompanied jacob allday’s expedition to rediscover greenland in 1759, and was noted especially for his diary of the voyage. krogen 70ø-427 (70°16.5´n 27°03.0´w). peninsula on the south coast of gåsefjord. named during the 1967–72 ggu scoresby sund expeditions by georg sawatzki for the hook-like shape (krogen = the hook). krogh-hytta 73ø (73°23.0´n 23°11.6´w). norwegian hunting hut on the south side of gauss halvø east of the river in aina dal, built by milne land korrid oren fig. 56. the eastern end of korridoren, the deep glacier-filled valley that crosses the centre of milne land, and that provided a low-level flying route for helicopters during the geological survey of greenland 1967–72 scoresby sund expeditions. 237 arktisk næringsdrift in october 1930. john giæver and otto johnsen, who built the hut, had got to know rolf von krogh during the summer, when he had undertaken hydrographical observations with the nsiu expedition. the name was intended to apply to the general location as well as the hut, but never acquired this usage. rolf von krogh [1872–1951] combined long service in the norwe gian navy with active arctic exploration, and took part in many expeditions to svalbard from 1924, and was in charge of survey work in the east greenland fjords from 1930 to 1933. the hut has also been known as aina dal hytten. (von krogh). krogh-johansens isbanke 76ø (76°37.0´n 20°50.7´w). shallow part of dove bugt between kap bjarne nielsen, the ne point of edvard ø, and bratskæret, where hundreds of icebergs derived from brede bræ lie stranded. so named during the 1932 gefion expedition, after v. krogh-johansen, a member of the committee of øst grøn landske fangstkompagni. søkort-arkiv (danish nautical charts archive) uses the form krogh-johansen isfjeldsbanke. kroghsundet 73ø (73°53.9´n 22°14.9´w). sound between jackson ø and home forland. the name was used on the 1932a nsiu map for the present gulmann sund, and commemorates rolf von krogh. see also krogh-hytta. krognesfjellet 73ø 74ø (74°00.6´n 21°37.0´w). name sometimes used by norwegian hunters for the mountain behind the hunting station krogness, equivalent to the present frebold bjerg. krogness 74ø (74°02.8´n 21°46.8´w). norwegian hunting station about 2 km sw of kap stosch, built by the foldvik expedition in 1926, and also commonly called kapp krogness. the station was manned from 1926 to 1930 and 1935 to 1937. it was named after ole andreas krogness [1886–1934], norwegian geophysicist and director of the geophysical institute at tromsø. he was responsible for the establishment of the first norwegian weather station at myggbukta in 1922. (krognæshytta.) krognessfjellet 72ø (72°57.1´n 23°16.0´w). mountain ridge on cen tral geographical society ø. the name was used on the nsiu maps of lacmann (1937), and commemorates o.a. krogness. see krogness. krogsø 73ø-544 (73°01.5´n 28°13.8´w). small lake at the corner where ptarmigan gletscher meets nordenskiöld gletscher, named by j.m. wordie’s 1929 expedition as corner lake (krog = corner). krokdalen 72ø (72°57.5´n 23°52.9´w). valley on west geographical society ø draining north into sofia sund. used only on nsiu maps (lacmann 1937), and named for the shape (krok = hook). kroken 76ø (76°19.0´n 20°48.3´w). norwegian hunting hut built in august 1933 for john giæver’s expedition on the west side of tvillingerne, dove bugt. it has also been called tvillinghytten and nordre jægersund hytten. krokfjellet 73ø (73°42´n 20°50´w). mountain 801 m high on east hold with hope. so named on the 1932a nsiu map for its hook-like shape. kronborg gletscher 69ø-32 (69°00.0´n 28°30´w). large glacier in east greenland. from the coast at c. 68°25´n it extends northwards to just beyond latitude 69°n. it was named after the noted castle, kronborg, in helsingør, denmark. kroneberghytten 74ø (74°34.9´n 19°13.5´w). danish hunting hut (so spelt) on the west side of sabine ø below kronebjerg. built by nanok in august 1948. kronebjerg 74ø-52 (74°35.4´n 19°08.8´w). mountain 544 m high on west sabine ø, named by karl koldewey’s 1869–70 expedition as kronenberg for the crown-like shape of the summit carved out of a basalt sill. sechsspitze and sevenspits have also been used. (kronenberge, mt kroneberg, landskrone, kronebjerget.) kronebjergpynten 74ø (74°34.9´n 19°13.5´w). name used by danish hunters for the locality beneath kronebjerg, sabine ø, where kroneberghytten was built in 1948. kronedal 71ø-409 (71°48.8´n 23°29.4´w). valley in the bjerg kronerne massive, draining south into ørsted dal. named by katherina perch-nielsen during the 1967–72 ggu scoresby sund expeditions. kronen 70ø-193 (70°35.5´n 22°06.8´w). snow-capped mountain in south liverpool land, named during the 1931–34 treårsekspedi tionen by laurits bruhn for its appearance (kronen = the crown). kronen 70ø-53 (70°42.9´n 25°28.0´w; map 4). prominent mountain 674 m high nw of kap leslie, east milne land. named by alfred rosenkrantz during lauge koch’s 1926–27 expeditions as mt kronen, for the shape of its summit (kronen = the crown). (kronenberg.) kronprins christian land 80ø-110 (80°45.0´n 20°00.0; maps 1, 4). extensive land area between danmark fjord to the west and the green land sea (grønlandshavet) to the east, with a southern bound ary at nioghalvfjerdsfjorden. so named by the 1906–08 danmarkekspeditionen after the then crown prince of denmark, later king christian x [1870–1947]. he became king in 1912. this region corresponds to the northern part of kong frederik viii land. kronprins frederik land 80ø (80°12.0´n 24°00.0´w). land area at the ne margin of the inland ice including a large part of north greenland west of kronprins christian land. it commemorates the journey made by kronprins frederik of denmark in 2000 as a member of the sirius sledge patrol. kronprins (crown prince) frederik [b. 1968] is heir to the danish throne. krosseyjar 76ø (c. 76°20´n 20°30´w). according to tornøe (1944) the group of islands in the sw part of dove bugt may have been the ‘krosseyjar’ of the icelandic sagas, but the identification is highly speculative. the description in bjørrn jónssons grønlands annaler is of four large islands surrounded by other small islands. if correctly identified the islands would represent the north limit of viking exploration in east greenland. (korsøy, krosseyjum, kaarsøø, kor søerne.) krumme langsø 73ø-92 74ø-334a (74°03.2´n 23°41.5´w; map 4). long lake in west hudson land, with a pronounced right-angled bend (krumme = bend). named by th. johansen during the 1931– 34 treårsekspeditionen. gaassjø has also been used. (krumme langsee.) krummedal 71ø-376 (71°24.0´n 29°00.0´w; map 4). valley with a pronounced hook-like shape, draining via rencontre dal to flyver fjord. named by peter vogt during lauge koch’s 1957 expedition (krumme = bend). krumodden 74ø (74°27.3´n 20°34.5´w). peninsula with a hookshaped termination on the coast of zackenberg bugt. the name has been used by scientists at zackenberg forskningsstation. krypt gletscher 72ø-496 (72°19.9´n 24°30.0´w). minor glacier in a deep, crypt-like valley in the syltopperne, north stauning alper. named by the 1963 university of cambridge expedition who as cended the glacier on their route to menander spir. (crypt glacier.) kuglelejet 81ø (81°13.0´n 13°52.3´w). area in central kilen, kron prins christian land, where a synclinal structure is developed in rocks containing football-sized concretions. the name is found on a coloured geological map of kilen printed in 1991 (pedersen 1991). kuhn ø 74ø-32 75ø-21a (74°50.4´n 20°15.3´w; maps 2, 4; fig. 15). large island nw of wollaston forland. named by karl koldewey’s 1869–70 expedition as kuhn insel, after the austrian war minister, baron franz kuhn [1817–1896], who supplied generous quantities of rifles and ammunition to the expedition. it is occasionally re ferred to in the expedition reports by the temporary name kohlen insel. (kuhn island, kuhnön, kunoøya.) kuhnpashytten 74ø (74°28.8´n 20°22.9´w). danish hunting hut west of kuhnpasset, wollaston forland, about 6 km from the coast. built by nanok in july 1951. (kuhn pas-hytten.) kuhnpasset 74ø-156 (74°29.9´n 20°20.6´w). pass in wollaston for land between cardiocerasbjerg and aucellabjerg. named during the 1931–34 treårsekspeditionen by hans frebold as kuhn pas, perhaps because it was used as a route to kuhn ø. 238 kûk, kûkajik, kûkajik kítikajik – see kuuk, kuukajik, kuukajik kitti kajik. kuldal 70ø-358 (70°08.5´n 22°13.0´w). small valley ne of the settlement at kap brewster, so named for a sequence of tertiary sediments containing three coal beds. greenlanders collected coal here. the name was used by hassan (1953) in his description of material collected during lauge koch’s 1951 expedition. kuldedal 71ø-309 (71°31.9´n 24°44.7´w). valley west of southern schuchert flod draining into ødemarksdal, eroded in barren, sterile sandstone. named by enrico kempter during lauge koch’s 1956–58 expeditions. kulfjeldet 75ø (75°11.5´n 19°59.8´w). coastal cliffs adjacent to kulhus, where danish hunters have mined substantial supplies of coal (hansen 1939; nyholm-poulsen 1985). the official name of this locality is jarners kulmine. (kulfjæld.) kulhus 75ø-62 (75°11.5´n 19°59.8´w). danish scientific station on the sw coast of hochstetter forland, built in 1932 during the 1931–34 treårsekspeditionen. it was named for the nearby outcrops of coal – see also kulfjeldet and jarners kulmine. the station was rarely used after the 1930s, and is now reported in poor condition (1990). (kulhuse.) kulhøj 77ø-25 (c. 77°26´n 21°33´w). hill at the nw end of annexssøen, so named by the 1906–08 danmark-ekspeditionen for the occurrence of abundant loose blocks of low-grade coal. the coal blocks are found over a wide area, and the exact location of the original finds is uncertain. (kulhöj.) kulisserna – see western upper terrace. kullabjerg 72ø-467 (72°50.3´n 28°58.5´w). nunatak in the upper part of nordenskiöld gletscher, sw of shackleton bjerg. the name derives from kullaberg (or kullen) in southern sweden, which features in the childrens story by selma lagerlöf (see also nils holgersen nunatakker). kullerne 70ø-418 (c. 70°58´n 27°56´w). moraine ridge 3 m high and 2 km long on the north side of harefjord. named during the 1967–72 ggu scoresby sund expeditions by svend funder. kumaat 70ø-295 (70°27.1´n 22°14.0´w). fossil locality at the coast on the west side of rosenvinge bugt, south liverpool land. the name was first recorded by alfred rosenkrantz in 1926, and reported to be that used by the settlers at kap hope [ittaajimmiit]. greenlandic inuit are said to refer to fossils as ‘the lice of the mountains’. (kumait.) kumaat 71ø-52 (71°42.8´n 22°35.4´w). mountain on the nw side of nathorst fjord, named by lauge koch’s 1926–27 expeditions as komait, after the inuit word for lice. numerous fossils were collected here. the original locality was at the south foot of the mountain which now bears the name. (kumait.) kumait – see kumaat. kumaqarteq 71-236a (c. 71°43´n 22°33´w). coastal stretch on the nw side of nathorst fjord. recorded by the 1955 geodætisk institut name registration, the name means ‘it has lice’ (= fossils) – see also kumaat. kuplen 70ø-426 (70°35.0´n 29°00.0´w; map 4). ice-capped mountain 1750 m high north of rolige bræ. named by laurent jemelin during the 1967–72 ggu scoresby sund expeditions for the dome-like appearance. kuplen 74ø-155 (74°23.3´n 20°09.3´w). mountain 506 m high in west wollaston forland, named during the 1931–34 treårs ekspedi tionen by hans frebold (kuplen = the dome). kuppelpashytten 74ø (74°24.3´n 20°04.9´w). danish hunting hut near kuppelpasset, near the head of isdal. built by nanok in may 1946. kuppelpasset 74ø-363 (74°24.3´n 20°06.1´w). pass at the head of isdal, west wollaston forland, north of the mountain kuplen. so named probably by nanok. the name appears on a map in jennov (1939) and also scientific reports from the same year. hunters used the pass as a route across wollaston forland. kursbræ 76ø-129 (76°33.0´n 24°48.0´w; map 4). glacier in south dronning louise land draining east into farimagdal. so named by j.p. koch’s 1912–13 expedition because by following its direction the expedition regained their planned route (kurs = bearing). (kurs bræen.) kuuk 70ø-307 (70°29.3´n 21°57.1´w). river flowing in elvdal, south ern liverpool land, entering the sea at ittoqqortoormiit [scoresbysund]. recorded by the 1955 geodætisk institut name registration, the name means ‘the river’. (kûk.) kuuk iserdoq 70ø (70°31.2´n 21°57.6´w). name reported by the scores bysund local newspaper in 1984 as in use by the inhabitants of the town for mågeelv. it translates as ‘muddy river’. kuukajik 70ø-311 (70°28.7´n 21°54.0´w). river in south liverpool land, flowing into amdrup havn. one of the names recorded by the 1955 geodætisk institut name registration, the name means ‘the little river’. (kûkajik.) kuukajik 71ø-225 (71°18.3´n 25°00.3´w). small stream ne of sydkap. the name was recorded by the 1955 geodætisk institut name registration, and translates as ‘the little river’. (kûkajik.) kuukajik 72ø-280 (72°56.2´n 25°21.9´w). river in ørkendal, se suess land. the name was recorded by the 1955 geodætisk insti tut name registration, and translates as ‘the little river’. (kûkajik.) kuukajik kittikajik [brudelv] 70ø-296 (70°30.0´n 22°13.2´w). river on the west side of rosenvinge bugt, southern liverpool land. recorded by the 1955 geodætisk institut name registration, the name translates as ‘the western little river’. (kûkajik kítikajik.) kvaksnes 72ø (72°43.5´n 22°51´w). sw peninsula of magga ø in vega sund. so named on the nsiu maps of lacmann (1937), be cause it resembles in shape the body of a wasp (kvaks = hveps = wasp). kvalen 73ø (73°06.4´n 22°27.5´w). island in the broch øer group. so named on the 1932a nsiu map, possibly from its appearance. the name may derive from the norwegian dialect word for a whale. kvalpen 72ø (72°44.5´n 22°45.5´w; fig. 14). small island in vega sund, east of the scott keltie øer. used only on nsiu maps (lac mann 1937), the name was given for its diminutive size (kvalp/ hvalp = whelp, young dog). kvarven 72ø (72°59.3´n 23°09.5´w). mountain on central geo graphical society ø, sw of rudbeck bjerg. so named on the nsiu (1932a) map, and in lacmann (1937), after a mountain of the same name in the nordland district of norway. kvelnadalen 74ø (74°10.1´n 20°19.1´w). valley on east clavering ø draining south into lervig. so named on the nsiu maps of lacmann (1937) after the norwegian river kvelna. kvina 73ø (73°48.0´n 21°42.0´w). river on the east coast of loch fyne, south of midtstua. so named on an nsiu map (1932a), possibly after a river of the same name in the vest-agda area of south norway. kvisladalen 74ø (74°13.7´n 21°34.1´w). valley on sw clavering ø, the present vildbækdalen. the name is used on the nsiu (1932a) map, and by lacmann (1937), and derives from the norwegian dialect word to describe the turbulent noise of a swift-flowing river. kvitegga 71ø (71°50.0´n 25°24.5´w; map 5). minor snow summit about 2400 m high on the south side of darien pass, just south of the snow peak darien, on the divide between the heads of bjørnbo gletscher (main glacier) and spærregletscher, stauning alper. it was climbed and so named by the 1996 norwegian stauning alper expedition. kvitfjell 72ø (72°04.0´n 24°52.0´w; map 5). mountain about 2350 m high on the ridge between the heads of gullygletscher and schu chert gletscher. climbed and so named by the 1996 norwegian stauning alper expedition (kvit = white). kyber 76ø (76°51.9´n 23°41.8´w). surveying station in central dron ning louise land. the name was used in the report on the 1952–54 british north greenland expedition by hamilton et al. 239 (1956), and was presumably named after the kyber pass. kystens perle 72ø (72°52.6´n 25°06.7´w). name by which the sirius summer station on ella ø, beside lauge koch’s scientific station, is affectionately known (bjerre 1980). it is also the name of a noted danish restaurant in kastrup, near copenhagen (café kystens perle). see also ella ø station. kystens perle 73ø (73°40´n 21°50´w). name by which danish trappers commonly referred to loch fyne station, in the inner part of loch fyne. kystfjæld 75ø (75°10.0´n 19°56.3´w). name used by danish hunters for part of søndre muschelbjerg in hochstetter forland, close to the coast (nyholm-poulsen 1985). kystkærene 74ø (74°27.6´n 20°32.5´w). boggy area along the coast of zackenberg bugt, south of zackenberg forskningsstation. the name is used as a reference locality in reports by visiting scientists. kaares-bu 71ø (c. 71°44´n 22°29´w). norwegian hunting hut in nathorst fjord, 6–7 km south of kap brown. built in august 1932 by helge ingstad and normann andersen, and named after ing stad’s brother, kaare [b. 1901], a diplomat who was norway’s ambassador in tel aviv from 1966 to 1971. the hut was moved to fleming fjord in 1955, where it is known as flemming fjord nord. (kåresbu, pass-huset.) kæmpebænken 73ø-572 (73°33.4´n 30°26.2´w). nunatak west of hamberg gletscher. named by arne høygaard and martin mehr en’s 1931 expedition as kjempebenken, and described as a strange box-like nunatak 800 m long, 400 m wide and 75 m high, resembling a giant bench. (kämperbänken.) kæmpegletscher 72ø-482 (72°18.1´n 26°17.3´w). broad glacier at the head of schaffhauserdalen, nathorst land. named by john haller following explorations during lauge koch’s geological expedition in 1954 (kæmpe = giant). kæmpehøjen 70ø-196 (70°37.0´n 22°01.5´w). snow-covered mountain 1050 m high in southern liverpool land, north of scores bysund. named by g.c. amdrup’s 1898–1900 expedition as kæmpehöjen for its large size. (kæmpehøj). kærelv 70ø-161 (70°47.3´n 22°26.6´w). small river in south liver pool land draining west into hurry inlet. named during the 1931–34 treårsekspeditionen by laurits bruhn (kær = marsh). kærdal 74ø (74°28.1´n 20°30.9´w). minor valley east of zackenberg forskningsstation in which kærelv flows. the name is used as a reference locality by visiting scientists. kærelv 74ø (74°28.1´n 20°30.9´w). minor river east of zackenberg forskningsstation draining south into young sund. the name is used as a reference locality by visiting scientists (meltofte & thing 1996). (fen river.) käthispids 72ø (72°11.0´n 24°59.6´w; map 5). minor peak 2350 m high at the head of vikingebræ climbed by peter braun and fritz schwarzenbach in late august 1950. it was named after a girlfriend of peter braun. kødgravene 80ø-57 (80°49.0´n 14°12.0´w; map 4). coastal area in ne amdrup land, north of sophus müllers næs. so named by the 1938–39 mørkefjord expedition because of the numerous stone mounds which proved to be inuit meat caches (kød = meat). (ködgravene.) l l. bistrup bræ 75ø-86a 76ø-109 (76°30.0´n 23°00.0´w; maps 2, 4; fig. 21). large glacier flowing northwards between dronning louise land and rechnitzer land. named by henning bistrup as l. bistrupsbræ during the 1906–08 danmark-ekspeditionen, after his father lauritz hans christian bistrup [1850–1914], who was a colony manager in west greenland. see also kap anna bistrup. (bistrups bræ, bistrup-bræ, l. bistrup glacier, bistrupsjökull.) l’acropole 71ø (71°55.8´n 25°57.5´w). nunatak at the head of glacier des oubliettes on the west side of prinsessegletscher. named and climbed by claude rey’s 1968 expedition. la cour bjerg 73ø-69 (73°31.1´n 22°32.8´w). mountain 1031 m high on gauss halvø, named by lauge koch’s 1929–30 expeditions in the form mt. la cour after dan barfod la cour [1876–1942], a physicist and meteorologist. he was director of the danish meteorological institute from 1923. la place huset, laplace – see laplace huset. la placeneset 73ø (73°00.5´n 22°30.8´w). peninsula on the north coast of geographical society ø, north of laplace bjerg. the name was used on an nsiu map (1932a), and was adopted by den grønlandske lods (1968) in the form laplacenæsset. labben 72ø-202 (72°13.8´n 23°48.9´w). peninsula west of the mouth of noret, north scoresby land. named by prospecting teams associated with lauge koch’s 1948–49 expeditions. labyrinth glacier 71ø (71°09.0´n 26°09.9´w). glacier on the south side of edward bailey gletscher, renland. named by the 2007 west lancashire mountaineering group expedition. lächen – see after lystergletscher (‘ä’ is treated as ‘æ’ in danish). lacmannfjellet 74ø (74°21.5´n 20°50.1´w). mountain on north clavering ø. so named on the nsiu maps of lacmann (1937), after otto lacmann [1887–1961], a pioneer of photogrammetry and professor at the technische hochschule berlin, who was involved in norwegian map-making in the arctic from 1919, and prepared the description of the three 1:100 000 scale norwegian maps of parts of east greenland. lacroix bjerg 73ø-625 (73°26.7´n 26°53.7´w; map 4). mountain with ice-capped summit about 2100 m high in sw andrée land, on the ne side of isfjord. named during the 1931–34 treårseks peditionen by eugène wegmann in the form lacroix mts; it was said to have been named after several french scientists, including the geologist françoise a.a. lacroix [1863–1948]. ladderbjerg 73ø-37 (73°35.2´n 22°09.6´w). northernmost peak of the giesecke bjerge on east gauss halvø, named by j.m. wordie’s 1926 expedition as ladder mountain, presumably for the step-like appearance caused by horizontal lava flows. the english form ‘ladder’ was retained in the approved name because it had been extensively used in publications. (ladderfjeldet.) lady øer 78ø-51 (77°59.0´n 20°26.7´w; map 4). small island group west of the danske øer. the name was one of a group of five given by the place name committee after dogs used on the 1906–08 danmark-ekspeditionen, that replaced names suggested by john haller. ‘lady’ was noted for running away with another dog on a sledge journey, and arriving home after six weeks’ absence in excellent condition. she was later found to have lived a life of luxury at a food depot. laffon bjerg 73ø-320 (73°48.5´n 23°14.7´w). mountain 1402 m high in hudson land, south of ritomsø. named by heinrich bütler during lauge koch’s 1936–38 expeditions after the naturalist johann conrad laffon [1801–82] of schaffhausen. (laffons bjerg, laffonberg.) lagerberghytte 72ø (72°31.2´n 24°39.5´w). norwegian hunting hut at kap lagerberg, se lyell land, built by the møre expedition in august 1930. it was originally known as beinhaugen. (kap lager berg hytten.) lagerholmen 74ø (74°30.0´n 18°57.0´w). name used by the 1908–09 floren expedition, probably for the small island off kap wynn which they also called maageholmen (brandal 1930). during the expedition they had deposited stocks of coal and salt here (lager = depot). lagernunatak 73ø (73°57.7´n 29°29.1´w). name used by katz (1952) for a nunatak west of orienteringsnunatak where he camped on 5 august 1951 (lager = depot). lagertoppen 71ø (71°49.2´n 25°39.2´w; map 5). peak in the ne part of the borgbjerg gletscher region, southern stauning alper. probably named by the 1977 schwäbische stauning alper expedition. 240 lagfjeld 73ø-97 (73°38.9´n 23°59.5´w; map 4). cliff up to 1821 m high on western gauss halvø, named during the 1931–34 treårs ekspeditionen by th. johansen for the striped appearance (lag = layer). the more common usage of the name is lagfjeldet. lagunen 74ø-308 (74°05.8´n 21°16.5´w). small lagoon that builds up behind the beach ridge of østhavn, adjacent to eskimonæs station. the name originated from the wintering party at eskimo næs during the 1931–34 treårsekspeditionen. lagunenæs 76ø-296 (76°55.0´n 20°13.5´w). minor cape between mørkefjord station and hvalrosodden. named by the 1938–39 mørkefjord expedition (lagune = lagoon). lagunenæsdal 71ø-396 (71°41.1´n 22°51.1´w). valley on wegener halvø reaching fleming fjord at lagunenæset. the name was adopted by katherina perch-nielsen during the 1967–72 ggu scoresby sund expeditions after the lagunenæsset valley of grasmück & trümpy (1969) and the lagunenæsset dal of trümpy (1969), names used during lauge koch’s 1958 expedition. see also lagunenæsset. lagunenæsset 71ø-84 (71°41.1´n 22°52.9´w). peninsula on the nw side of wegener halvø, so named during the 1931–34 treårs ekspeditionen by arne noe-nygaard for a coastal lagoon. (lagu ne næs.) lake b1 72ø (c. 72°42´n 22°29´w). lake on geographical society ø where samples were collected for radiocarbon age determinations (cremer et al. 2008). lake n1 73ø (c. 73°20´n 25°13´w). lake on ymer ø where samples were collected for radiocarbon age determinations (cremer et al. 2008). lakse sø 77ø (77°04.5´n 20°50.4´w). name sometimes used in reports of the 1906–08 danmark-ekspeditionen for sælsøen, from which lakseelven drains. the arctic char found in rivers and lakes in east greenland is commonly referred to as salmon (= laks). lakseelv 70ø-280 (70°30.0´n 22°42.1´w; map 4). river in se jameson land west of kap stewart. named by alfred rosenkrantz during lauge koch’s 1926–27 expeditions as lakse elv after the arctic char (fig. 57; laks = salmon). lakseelv 73ø-604 (73°43.3´n 24°35.8´w; map 4). river draining laksesø in brogetdal, east strindberg land. salmon (= arctic char) were regularly fished here by norwegian hunters. (salmon river.) see also giæver-hytta. lakseelva 74ø (74°27.5´n 21°41.1´w). norwegian hunters name for a river in tyrolerfjord where they fished in the summer, probably that east of giesecke bjerg. see also giskehuset. lakseelven 74ø-183 (74°51.0´n 20°53.1´w). river draining blåbærdalen, east th. thomsen land, draining into fligely fjord. the name is said to have been given by danish hunters in 1929, and first appeared in print on the 1932 edition of the geodætisk institut 1:1 million scale map prepared during the 1931–34 treårs ekspeditionen. (lakselv, lakseelv.) lakseelven 75ø (75°15.0´n 20°57.4´w). name occasionally used by danish hunters for the river draining kildedal on the south side of ardencaple fjord, which they also called kildeelven (nyholmpoulsen 1985). lakseelven 76ø-61 (76°55.5´n 20°09.1´w). river draining sælsøen. so named by the 1906–08 danmark-ekspeditionen because of the many salmon (arctic char), of which 200 kg were caught here in august 1907. (salmon river, laxá.) laksehytta 73ø (73°42.2´n 24°30.6´w). norwegian summer station at the mouth of brogetdal, strindberg land. it was built by arktisk næringsdrift in june 1935 for fishing, and attempts at packing the ‘salmon’ (arctic char) in tins were made here in 1938. it is also known as strindberghuset. laksehytta 73ø (73°19.1´n 25°02.8´w). hut at the head of dusén fjord, ymer ø, built in august 1932 by the crew of the isbjørn to support fishing. it is also known as noahytten, bunnhuset and holmboe-hytta. laksehytten 74ø (74°27.9´n 20°39.1´w). norwegian hut built for salmon fishing in the summer of 1949 west of zackenberg hunting station for herman andresen’s expeditions. it is also known as fiskerhytten. laksehytten 76ø-209 (c. 76°07´n 20°29´w). danish hunting hut on the ne shore of laksesø, ad. s. jensen land, said to have been built by nanok in 1939. although this name is officially approved, the hut was never built (p.s. mikkelsen 1994). it may have been confused with the danish hut fiskerhytten between syttendemaj fjorden and laksesø. laksesø 72ø-229 (72°07.9´n 23°42.9´w). small lake on the east side of mesters vig. named by prospecting teams associated with lauge koch’s 1948–49 expedition. laksesø 73ø-366 (73°43.7´n 24°40.4´w). largest of the lakes in brogetdal, strindberg land. named during lauge koch’s 1948–49 expeditions by hans r. katz, after the abundant arctic char. laksesø 76ø-191 (76°10.2´n 20°41.9´w). lake in ad. s. jensen land at the head of syttendemajfjorden. the name was proposed by nanok, and appears on a map in jennov (1939). lambert land 79ø-1 (79°15.0´n 20°40.0´w; maps 1, 4). land area almost surrounded by the glaciers of nioghalvfjerdsfjorden and zachariae isstrøm. adapted by the 1906–08 danmark-ekspedi tionen from an old dutch chart from 1718, which reported t'land van lambert to have been discovered at this latitude by a whaler of that name in 1670. (lamberts land.) lambeth 72ø (72°05.8´n 24°54.9´w; map 5). mountain 2450 m high between gully gletscher and the head of bersærkerbræ, north stauning alper. first climbed by the 1963 imperial college expedition, and named after the south london borough. lamorna 72ø (72°08.1´n 24°55.2´w; map 5). pinnacle about 2700 m high on the ne ridge of hjørnespids, north stauning alper. named and climbed by the queen mary college expedition on 13 august 1968. lamprenen dal 71ø-177 (71°38.0´n 23°38.8´w; map 4). valley west of fleming fjord draining nw into ørsted dal. the name was one of a group of names given by the place name committee in 1939 to replace proposals by hans stauber. it was given for one of the ships used by jens munk on his voyage in search of the northwest passage in 1619. (lamprenens dal.) lancaster 71ø (71°46.6´n 25°32.9´w; map 5). peak about 2510 m high in the south stauning alper between borgbjerg gletscher and orion gletscher. climbed by the 1971 university of lancaster expedition. see also lancaster bugt. lancaster bugt 71ø-440 (71°35.0´n 27°58.0´w; map 4). deep bay on the north side of flyverfjord. named by geoffrey halliday during the 1971 northern universities expedition after the university of lancaster, to which he was affiliated. the town of lancaster grew up on the site of a roman fortification, while the university was founded in 1964. garagebugt has also been used. landhuset 71ø (71°33.1´n 22°58.1´w). norwegian hunting hut built in 1932 or 1933 for helge ingstad’s expedition in pingel dal, about 12 km south of the head of fleming dal. it is also known as fleming dal hytten and pingel dal hytten. landingsdalen 74ø-187 (74°27.5´n 19°03.1´w). valley in east wollaston forland, south of kap wynn. so named by nsiu in 1929 when the veslekari was unable to reach the huts at kap wynn due to ice conditions, and landed all their supplies at the mouth of this valley. landtungen 71ø (71°20.6´n 24°36.9´w). name used by kempter (1961) for the tongue of land between nordostbugt and schuchert flod. langtungen was used by the 1962 oxford university expedion for the same feature (sugden & john 1965). lang peak 1, 2, 3, 4, 5, 6 71ø (c. 71°56´n 24°36´w to 71°59´n 21°44´w; map 5). series of six summits ranging from 1940 m to 2100 m in altitude on the ridge ne of storgletscher, central stauning alper. storgletscher was for a period known as langgletscher. the 1961 241 bangor mountaineering club expedition named the peaks, and climbed numbers 2, 3, 5 and 6. langbjerg 72ø-256 (72°55.8´n 22°42.9´w). elongate mountain ridge up to 522 m high on geographical society ø, so named during lauge koch’s 1949–50 expedition by desmond t. dono van. langbjerg 73ø-289 (73°30.5´n 22°49.6´w). elongate n–s ridge on gauss halvø, named during the 1931–34 treårsekspeditionen by gunnar säve-söderbergh as long mountain. langdyssen 72ø-200 (72°14.4´n 23°55.0´w). small ridge west of noret, north scoresby land, resembling an elongate burial mound (langdyssen = long barrow). named by prospecting teams associated with lauge koch’s 1948–49 expeditions. langdyssen pools 72ø (c. 72°14´n 23°55´w). name used by the 1968–74 university of dundee expeditions for five interconnected pools near langdyssen at the ne end of mestersvig airfield. langefirn 71ø-259 (71°59.0´n 24°09.8´w; map 5). glacier in the werner bjerge flowing west to join arcturus gletscher. named during the 1953–54 lauge koch expeditions by peter bearth and eduard wenk. langelandselv 70ø-101 (70°34.0´n 23°22.8´w; map 4). river in southern jameson land flowing south into scoresby sund. named during the 1931–34 treårsekspeditionen by laurits bruhn after the island of langeland, denmark. langelinie 72ø-189 (72°09.1´n 24°06.9´w; map 5). mountain ridge rising to 1058 m south of store blydal, north scoresby land. named by prospecting teams associated with lauge koch’s 1948–49 expeditions after the langelinie in copenhagen harbour between kastellet and frihavnen, where the little mermaid is located. langelinie 74ø-111 (74°15.3´n 20°31.6´w). mountain ridge about 800 m high on east clavering ø, named by lauge koch’s 1929–30 expeditions in the form mt. langelinie after the locality of the same name in copenhagen, denmark. (langelinie bjærg.) langelv 75ø-64 (75°44.2´n 20°00.0´w; map 4). river draining langsø and knæksø in the interior of nørlund land, entering rose neathbugt on the south side of mønstedhus. the name is attributed to the wintering party at kulhus in 1935, and first appears on a map in jennov (1939). langelv-hytten 75ø (c. 75°45´n 20°03´w). norwegian hunting hut about 15 km from the mouth of langelv on the right bank, built by arktisk næringgsdrift in 1932, and rebuilt in 1950. langelv fiskerhytte 75ø (75°41.9´n 19°34.1´w). norwegian hut built for salmon fishing by arktisk næringsdrift in june 1949 on the south side of langelv, about 500 m south of mønstedhus. langemands sø 74ø (74°30.1´n 20°36.2´w). small lake in the area known as morænebakkerne, north of zackenberg forsknings station. the name is used as a reference locality by scientists studying lake ecosystems. (langemandssø, langemand sø.) langemanden 70ø-439 (70°33.2´n 29°19.2´w). glacier between rolige bræ and døde bræ. named by laurent jemelin during the 1967–72 ggu scoresby sund expeditions for its length and the association with djævlehånden (langemand = long finger). langenthaler col 72ø (72°43.5´n 27°30.8´w). broad, flat col at the head of langenthaler gletscher leading to the nw part of glet scher land. the name was used by the 2002 shackleton bjerg expedition, which used this route to reach the ice cap and climb shackleton bjerg. langenthaler gletscher 72ø-462 (72°46.9´n 27°20.3´w). glacier in north gletscherland, draining north to the head of dickson fjord. the name was used by eugène wegmann during the 1931– 34 treårsekspeditionen, and was given for langenthal, a municipality in the canton of bern, switzerland. the glacier was ascended by eugène wegmann, augusto gansser and others on 10 august 1933 during their exploration of inner gletscherland. (langdals gletscheren, langdal bræ.) langenæs 70ø-9 (70°34.0´n 28°13.5´w; map 4). long, narrow peninsula between the front of rolige bræ and vestfjord, named by carl ryder’s 1891–92 expedition. a party from the expedition camped here during their first sledge journey in april 1892. (lange næs.) langeryg 73ø-411 (73°59.8´n 28°30.8´w). elongate nunatak in arnold escher land, named by hans r. katz during lauge koch’s 1951 expedition. langesø 70ø-60 (70°29.7´n 26°13.0´w). the largest lake on danmark ø, named during carl ryder’s 1891–92 expedition. langevåg 73ø-n287 (73°33.5´n 20°27.0´w). wide bay on the east coast of hold with hope, sw of holland ø. so named on the 1932a nsiu map because of the size of the bay. langgletscher 71ø (71°57.0´n 24°43.0´w). name occasionally used, and also briefly officially approved, for the long glacier in the stauning alper flowing east to schuchert dal, now known as storgletscher. bjørn jorsalfarers gletscher has also been used, and was the approved name between 1939 and 1971. langgletscher 72ø-153 (72°20.1´n 22°37.3´w). glacier on se traill ø south of mountnorris fjord. named during lauge koch’s 1936– 38 expeditions by hans p. schaub, for its length relative to other glaciers. (lang-gletscher.) langgletscher dal 73ø-433 (73°04.1´n 25°38.7´w). valley in ne fig. 57. laks (salmon) is the name commonly used for the arctic char in northern east greenland. they were formerly caught in large numbers using nets at the mouths of rivers. in the national park area at the present day fishing is only permitted using rods. 242 suess land that drains north into antarctic sund, named for the long glacier that occupies most of the valley. named during lauge koch’s 1947–49 expeditions by silvio eha. langholmen 77ø-63 (77°25.8´n 20°10.9´w). island in the inner part of skærfjorden, so named during the 1931–34 treårsekspe ditionen by david malmquist. it was named after the locality långholmen in central stockholm, sweden. on malmquist’s original maps it is elongate in shape, but on modern maps is almost circular. langkløftgletscher 72ø-477 (72°05.2´n 26°16.3´w). glacier north of furesø, north of castorbjerg. so named during lauge koch’s 1954–55 expeditions by hans zweifel, because the glacier occupies a long, narrow ravine. langryggen 70ø-263 (70°03.0´n 24°00.0´w). n–s-trending ridge up to 1800 m high on the west side of vestre borggletscher, south of scoresby sund. so named during the 1931–34 treårsekspe ditionen by laurits bruhn. langsiden 73ø-129 (73°48.9´n 20°35.4´w). n–s-trending ridge in home forland. named originally in the form langsuen on the nsiu (1932a) map. langsø 72ø-125 (72°52.5´n 25°07.5´w). elongate lake on nw ella ø, so named by the 1931–32 wintering party on ella ø. (langesø, long lake, langsee.) langsø 75ø-43 (75°48.9´n 20°48.0´w; map 4). long lake in dronning margrethe ii land, that first appears with this name on the 1932 geodætisk institut 1:1 million scale map compiled by lauge koch. langsødalen 75ø-110 (75°49.9´n 20°49.0´w). very long e–wtrending valley in hochstetter forland that contains the lake langsø. although commonly used, this name was not officially approved until 1981. langsøhytten 75ø (c. 75°42´n 19°33´w). danish hunting hut built by nanok in august 1933 on the north side of langelv. it was rescued from falling into the sea by j.g. jennov in 1948, who moved it nine metres inland. it functioned as a provisions shed after mønstedhus was built on the same site in 1938. following severe coastal erosion langsøhytten was taken by the sea in 2001, and mønstedhus was lost in the same way in september 2002. langtungen – see landtungen. langåren 73ø-246 (73°04.3´n 22°41.6´w). elongate island in the brochs øer group, first named on the 1932a nsiu map in the form langåra. (langaaren.) laplace bjerg 72ø-13 (72°5.16´n 22°32.7´w; map 4; fig. 12). moun tain 1190 m high on geographical society ø. william scores by jr. in 1822 gave the name cape laplace out of respect to pierre simon, marquis de laplace [1749–1827], a mathematician and astronomer. both spellings of his name (laplace and la place) are found in biographies and scoresby’s (1823) narrative. white (1927) observed that scoresby’s cape was easily identifiable with a mountain on geographical society ø which he called mount la place, and that has become laplace bjerg. (la placefjellet, la placefj.) laplace huset 73ø (73°00.0´n 22°31.9´w). name sometimes used for the norwegian hunting station built in 1938 at the foot of laplace bjerg, geographical society ø, by ole klokset’s expedition. it was manned as a wintering station only in 1938–39, and is now in poor condition. (laplace, kap laplace, la place huset.) laplace øer 72ø (73°00.7´n 22°30.3´w). low islands ne of geo graphical society ø. the name is used in den grønlandske lods (1968), and they are probably the small islands off laplaceneset. laplaceneset 73ø (73°00.8´n 22°30.6´w). cape on the north coast of geographical society ø, due north of laplace bjerg. the name is used on lacmann’s (1937) maps, and also in den grønlandske lods (1968). lapstun hytten 71ø (71°52.2´n 22°45.6´w). norwegian hunting hut erected in september 1954 by otto lapstun for herman andre sen’s expedition on the nw side of fleming fjord. it replaced the nearby hut flatstranda, which was swept away by a storm in 1953. the new hut has also sometimes been known as flatstranda, as well as surøje-hytten, søndre biot and fleming fjord hytten. large débris cone – see western upper terrace. lars christensenfonna 74ø (74°16.2´n 21°12.7´w). ice cap on central clavering ø, the present snemarken. so named on the nsiu maps of lacmann (1937) after consul lars christensen [1884–1965], a norwegian whaling magnate, ship owner and philanthropist. he made major contributions to antarctic exploration, subsidising a series of antarctic expeditions between 1926 and 1937. he also subsidised norwegian activities in east greenland, notably the re-establishment of myggbukta radio station in 1930, and the loan of an aeroplane for the 1932 aerial photography. lars jakobsen pynt 74ø-84 (74°32.4´n 19°10.8´w). peninsula on the south side of heimland havn, west sabine ø. named by j.m. wordie’s 1926 expedition as lars jakobsen point after the captain of the heimland, the expedition ship. larsens skær 76ø (c. 75°25.5´n 20°07´w). skerry 5 km ne of teufelkap in dove bugt. discovered during the 1932 gefion expedition (jennov 1935), and named after a business acquaintance named larsen. the name is used in den grønlandske lods (1968). larsfjeldene 77ø-67 (77°31.0´n 20°32.0´w). mountains in east nordmarken between agsutsund and v. clausen fjord, on both sides of h.g. backlund fjord, so named during the 1931–34 treårsekspeditionen by david malmquist. the name has been changed in recent official name lists to larsfjeld. laub nunatakker 78ø-18a (78°03.0´n 23°00.0´w; map 4). nunatak group west of hertugen af orléans land. named by the 1909–12 alabama expedition after vilhelm laub, second in command of the expedition, who led a party to the west side of dronning louise land. vilhelm laub [1887–1945] later became a director of the østasiatisk kompagni, and in 1932 director of copenhagen harbour. (laub’s nunatakker, laub’s nunataks.) lauge koch bjerg 72ø-484 73ø-557 (72°59.7´n 27°57.2´w; map 4). mountain 2436 m high in goodenough land. j.m. wordie gave it the name mount lauge koch in 1929, as a mark of his respect for koch’s work. lauge koch [1892–1964], a danish geologist and greenland explorer, took part in the 1916–18 second thule expedition, led the 1920–23 jubilæumsekspedition nord om grønland, and became most noted for his long series of geological expeditions to east greenland between 1926 and 1958. (lauge kochs bjerg.) lauge koch vig 75ø-57 (75°17´n 20°04´w; map 4). bay on the sw coast of hochstetter forland. the name appears to have first been used on a sketch map made by t. johansen in 1932 during the 1931–34 treårsekspeditionen, and was subsequently used by koch (1940). see also lauge koch bjerg. (lauge koch’s vig.) laugeites ravine 74ø (74°47.4´n 20°33.9´w). small ravine on se kuhn ø, named by maync (1947) for finds of fossils during lauge koch’s 1936–38 expeditions. the name, although unapproved, has subsequently been used as a type locality for a geological unit. laupdalen 74ø (74°24.5´n 20°45.0´w). valley on north clavering ø. used on the nsiu maps of lacmann (1937), the name is derived from a norwegian dialect word (laupa = shine, glisten). laussedatfjellet 74ø (74°20.7´n 21°07.9´w). mountain 1525 m high on north clavering ø. used only on nsiu maps (lacmann 1937), the name was given for aimé laussedat [1819–1907], a french officer and scientist credited with the invention of photogrammetry. he became professor of geodesy at the école polytechnique, and was subsequently director of the ‘conservatoire des arts et métieres’ in paris (national conservatory of arts and crafts). lavenesset 72ø (72°01.1´n 23°03.7´w). low peninsula on the se side of antarctic havn. the name was used by norwegian hunters, e.g. by jonas karlsbak in the excerpt from his 1930 diary published in tornøe (1944). it has also been called knivodden. lavinegletscher 73ø-612 (73°12.2´n 28°01.3´w). glacier on the 243 south side of knækdalen, named by louise boyd in 1933 as avalanche glacier because of the periodic ice-falls from the front (lavine = avalanche). lavkæret 74ø (74°28.0n 20°38.7´w). reference locality used by visitors to zackenberg forskningsstation for a boggy area north of zackenberg hunting station. lavøira 72ø (c. 72°13´n 23°45´w). norwegian hunting hut built in august 1930 for the møre expedition on the east side of hovedet, near mestersvig. it was also known as solstrand. it was moved in 1954 to fleming fjord. (lavøyra, laag-øyra, lavøren, lavøyra huset.) le casque 71ø (71°50.0´n 25°41.0´w; map 5). peak about 2450 m high at the head of prinsessegletscher, between col de scoresby and col de furesoe. named and first climbed by claude rey’s 1968 expedition. le nez blanc 70ø (70°46.4´n 21°59.7´w). minor summit in liver pool land, 892 m high, north of bjerring pedersen gletscher. climbed and named by the 2002 loughborough grammar school expedition. ledesia bjerg 74ø-371 (74°45.6´n 22°47.7´w; map 4). mountain or nunatak in the nw upper part of pasterze. named by the 1948 leeds university expedition, the name derives from the latin name for leeds. leeds bugt 71ø-441 (71°37.2´n 27°40.4´w; map 4). bay on the north side of flyverfjord. named by geoffrey halliday during the 1971 northern universities expedition after the university of leeds, to which the senior geologist david rex was affiliated. leeds in west yorkshire originated as an anglo-saxon settlement, while its university was founded in 1904. leicester bugt 71ø-377 (71°57.0´n 27°57.8´w; map 4). bay on the north side of inner nordvestfjord, east of backlund bjerg. named by geoffrey halliday during the 1961 leicester university expedition, which reached this point during botanical journeys. the university in leicester was founded as a college in 1918, and re ceived its charter in 1957. leira 73ø (73°28.4´n 21°40.2´w). river flowing into mackenzie bugt. so named on an nsiu map (1932a; fig. 13) because of its clayey nature. (leira river.) leirasund 73ø (73°57.6´n 21°10.8´w). sound between stille ø and the delta at the mouth of gulelv. so named on the 1932a nsiu map because of the muddy water. leirdalen 73ø (73°01.7´n 23°11.9´w). valley on north geographical society ø, draining north into sofia sund, west of robertson ø. so named on nsiu maps of lacmann (1937) for the clay deposits. leirneset 73ø (73°00.7´n 23°06.7´w). clayey delta on the north side of geographical society ø, at the mouth of leirdalen. so named on the nsiu maps of lacmann (1937) for the clay. leirvågen 74ø (74°26.1´n 20°56.1´w). norwegian hunting hut built in august 1939 on the west side of lerbugt, originally for glacial investigations by hans w:son ahlmann. it was also known as eigerhytta. norwegians also used the name leirvågen for the bay. (lerbugt, ler bay, leirevaagen, lerbugthytta.) leitch bjerg 72ø-11 (72°51.0´n 22°27.8´w; map 4; fig. 12). mountain 726 m high on geographical society ø. william scores by jr. in 1822 gave the name as cape leitch, but his cape was recognised as a mountain by white (1927) and renamed mount leitch. nsiu maps placed their leitchfjellet a few kilometres to the wsw, and gave the name brandegga to the present leitch bjerg (lacmann 1937). rund top was used apparently for the same feature by carl ryder’s 1891–92 expedition. (leitchfjellet.) lejrelv 71ø-56 (71°10.1´n 22°39.1´w). river in east jameson land, draining into klitdal. named during lauge koch’s 1926–27 expeditions by alfred rosenkrantz and tom harris, and used in various forms, lejr river, camp river or lejr elv, because they had a camp here. lejrgletscher 73ø-618 (73°13.9´n 27°59.7´w). glacier on the north side of knækdalen, named by louise boyd in 1933 as camp 2 glacier, because their second camp in the valley was sited close to the glacier front. (camp 2 gletscher.) lejrryggen 71ø-247 (71°58.5´n 23°56.1´w; map 5). ridge on the east side of østre gletscher, werner bjerge. so named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk, because bearth had a camp at the foot of the ridge during geological exploration in 1953. lembcke bjerg 77ø-48 (77°09.4´n 24°47.4´w; fig. 21). nunatak in nw dronning louise land, named during the 1909–12 alabama expedition as lembcke’s nunatak. preben lembcke [1886–1965] was a danish naval officer and a contemporary of wilhelm laub who had explored this area. lemenkjeften 73ø (73°27.2´n 21°36.5´w). small, enclosed bay on the west side of mackenzie bugt. named in this form on an nsiu map (1932a), after the lemmings. lemming bay 72ø (72°16.1´n 24°02.3´w). name used by the 1968–74 dundee university expeditions for the bay east of the mouth of skeldal elv. lemming lake 76ø (76°25.7´n 18°47.7´w). lake on store koldewey where sampling was undertaken for phytoplankton studies (cremer et al. 2005). lemming valley 72ø (72°15.0´n 24°01.0´w). name used by the 1968–74 dundee university expeditions for the broad valley west of mestersvig airfield. lemmingbugt 72ø-130 (72°51.9´n 24°54.4´w; map 4). bay on eastern ella ø, named during the 1931–34 treårsekspeditionen by the ella ø wintering party after the lemmings. lemmingdal 72ø-275 (72°51.9´n 24°57.1´w). valley on eastern ella ø, draining into lemmingbugt. named by john w. cowie during work carried out during lauge koch’s 1949–54 expeditions. lemmingelv 72ø-181 (72°51.9´n 24°57.1´w). name used for the river in lemmingdal, eastern ella ø, by christian poulsen about 1950. lemminghoved 73ø-670 (73°40.1´n 27°10.5´w). mountain ridge 1527 m high in west andrée land, at the east border of gerard de geer gletscher. so named by john haller following explorations during lauge koch’s 1950 expedition, because the rounded, long mountain ridge resembled in shape the head of a lemming. it was climbed by haller in 1950. (lemminghovet.) lenggrieser ryggen 71ø (71°49.7´n 24°58.2´w; map 5). mountain ridge about 2550 m high on the south side of the upper basin of spærregletscher. climbed by karl m. herligkoffer’s 1966 expedition on 19 august. lennox spids 72ø-355 (72°01.7´n 25°19.3´w; map 5). peak about 1800 m high on the sw side of sefström gletscher, north stauning alper. first climbed by malcolm slesser’s 1958 expedition, and named after the castle of lennoxlove, east lothian, scotland. (lennox.) leo gletscher 71ø-340 (71°38.8´n 24°55.2´w; map 5). glacier on the south side of bjørnbo gletscher, south stauning alper. so named by john hunt’s 1960 expedition after the constellation, the 5th sign of the zodiac. (leo glacier.) lepidoceraselv 71ø-74 (71°16.6´n 22°34.6´w). river in eastern jameson land draining east into carlsberg fjord. named during lauge koch’s 1926–27 expeditions by alfred rosenkrantz and tom harris as r. lepidoceras elv, for the fossil flowering plants. lepidurus loch 72ø (72°15.7´n 23°57.7´w). name used by the 1968–74 university of dundee expeditions for a very small lake in the hills west of nyhavn, because of the occurrence of ‘lepidurus arcticus’, a common freshwater entomostracean (daphnia). lerbugt 74ø-316 (74°25.8´n 20°55.5´w; map 4). bay on the north coast of clavering ø, named on nsiu 1937 maps in the form leirevag, for the clay deposits. (ler bugt, ler bay, leirevaagen, leirvågen.) lerbugthytta – see leirvågen. 244 lerelv 70ø-7 (70°45.2´n 28°59.2´w; map 4). large river draining into the west side of rødefjord. so named by carl ryder’s 1891–92 expedition because of the large banks of clay and silt formed at its mouth. lerelv 77ø-78 (77°19´n 19°57´w). clayey river draining into the south side of c.f. mourier fjord, west of kap li. the name was first used by david malmquist, following his surveying in the region with the 1931–34 treårsekspeditionen. lersøen 75ø (c.75°08´n 19°45´w). name used by danish hunters for a lake behind the hunting station nanok in southern hochstetter forland (hansen 1939). lervig 74ø-190 (74°09.4´n 20°20.2´w). small bay on the se coast of clavering ø. the name was first used by gelting (1934) as a botanical reference locality during the 1931–34 treårsekspeditionen, and records the clayey nature of the bay. lerø 76ø-253 (76°46.1´n 18°39.2´w). small island in østerelven, north of danmark havn, largely made up of clay (= ler). so named by the 1906–08 danmark-ekspeditionen, and first used in the description of the vegetation by lundager (1912). (lerö.) les cinq doigts 70ø (70°42.9´n 25°55.9´w). line of five nunataks on the south side of charcot gletscher, se milne land. so named by parat & drach (1934) in their report on j.b. charcot’s 1933 expedition. there were five in 1933, but three have merged, and there are now two nunataks with a long thin strip of land joined to the mainland to the east. les diablarets 73ø (73°12.9´n 27°48.1´w). name used in a climbing report by buess (1953) for the ridge on the north side of knæk dalen, opposite portgletscher. it was climbed by a party during explorations on lauge koch’s 1951 expedition, and apparently resembles the 3246 m peak above the swiss winter sports centre at les diablarets. leutkircher tinde 71ø (71°49.7´n 25°16.2´w; map 5). mountain on the sw side of roslin gletscher. climbed by karl m. herlig koffer’s 1966 expedition on 21 august, and named after the ba varian town of leutkirch. (leutkirchnertinde.) levynitfjeld 69ø-80 (69°54.6´n 27°18.8´w). mountain ridge 1826 m high between magga dan gletscher and kista dan gletscher. named by w. stuart watt during the 1967–72 ggu scoresby sund expeditions for an occurrence of the zeolite mineral levynit. liaselv 71ø-182 (71°19.6´n 22°38.6´w). river on the west side of carlsberg fjord north of liasryggen. so named during lauge koch’s 1936–38 expeditions by hans stauber because it cuts through rocks of liassic age. liasryggen 71ø-349 (71°18.1´n 22°39.2´w). mountain ridge 840 m high in eastern jameson land, west of inner carlsberg fjord. named in geological reports on work during the lauge koch expeditions by john h. callomon, for the age of the rocks. liavaag 74ø (74°29.4´n 18°59.6´w). norwegian hunting station built in 1929 for arktisk næringsdrift beside the 1928 gåsneshuset, about 1 km north of kap wynn. named after severin gaasnes liavaag, see also gåsneshuset. the two huts have also been known for their location as kap wynn hytter. (liavåg ), lichenryg 81ø (81°18.0´n 14°10.2´w). ridge in nw kilen, kronprins christian land, named after the black lichen that decorate the light-coloured sandstone slabs. the name is found on a coloured geological map of kilen printed in 1991 (pedersen 1991). licht ø 76ø-22 (76°27.4´n 20°25.5´w; map 4). island in sw dove bugt, so named by the 1906–08 danmark-ekspeditionen. possibly given for mathias kjeldsen de fine licht [1859–1917], a lieutenant in the danish navy and chief pilot (j. løve, personal communication 2009). (lichts ö, lichts island, lichts ø.) lidskjalv 74ø-295 (74°51´n 21°13´w). mountain about 1300 m high in th. thomsen land, east of odin dal. the name originated from the wintering party at kulhus during the 1931–34 tre årsekspeditionen, and was given for odin’s throne in nordic mythology. lilienthalflya 73ø (73°00.9´n 22°51.1´w). lower slopes of northern geographical society ø, wsw of tveholmen (flya = plain). the name is used only on nsiu maps (lacmann 1937), and commemorates otto lilienthal [1848–96], a german pioneer of gliding. his death in a gliding accident was reported to be the first ever arising from pilot error. lille blydal 72ø-211 (72°10.4´n 23°52.7´w). valley in northern scoresby land draining northwards into noret. it is separated from store blydal by a low pass. named by prospecting teams associated with lauge koch’s 1948–49 expeditions. lille bælt 76ø-74 (76°41´n 18°48´w; maps 2, 4). sound between lille koldewey and store koldewey. named by the 1906–08 danmark-ekspeditionen, and first used in the hydrographical report by trolle (1913). named after the sound of the same name in denmark between the island of fyn and jylland. see also store bælt. (lille-bælt, lille belt, lillebælt.) lille cervin 73ø-663 (73°25.8´n 27°36.3´w; map 4; fig. 58). moun tain about 1600 m high in northern frænkel land on the south side of jættegletscher. the name apparently arose independently from two sources, laurits bruhn of the geodætisk institut and noel e. odell. both remarked on its resemblance to the matterhorn (= monte cervino or mont cervin) on the border between switzerland and italy. the 1972 university of dundee expedition made an attempt to climb it, but did not reach the summit. (matterhorn south peak.) lille cirkusbjerg 71ø-107 (71°37.9´n 22°53.7´w). mountain on south wegener halvø, named during lauge koch’s 1936–38 expeditions by gunnar säve-söderbergh (1937) as little circus mountain. lille finsch ø 74ø (74°00.3´n 21°07.0´w). next largest of the finsch øer. distinguished first on maps of lauge koch’s 1929–30 expeditions as little finsch island, and on norwegian charts as vesle finsch; the name was briefly in use as an approved name, but later discontinued. lille kløft 70ø (70°39.9´n 22°40.9´w). small ravine on the west side of hurry inlet, north of moskusoksekløft. the name was used in a report by heinrich aldinger (1935) on work during the 1931–34 treårsekspeditionen. lille koldewey 76ø-38 (76°39.0´n 18°40.9´w; maps 2, 4). two islands separated by a narrow sound, røseløbet, situated ne of store koldewey. so named by the 1906–08 danmark-ekspedi tionen. north koldewey island was used by amdrup (1913) for the same feature. a depot left on the east coast of the northern island by the german meteorological ‘edelweiss’ expedition in 1944 is known as tyskedepot. (little koldewey, lilla-koldewey.) lille myteklippe 70ø-379 (70°15.1´n 29°00.9´w). cliff on the south coast of kaskadesø, western gåseland. so named during lauge koch’s 1958 expedition by eduard wenk, because it and the adjacent cliff (store myteklippe) were in their shape and tectonic relationships similar to the grossen mythen and kleinen mythen in canton schwyz, switzerland. lille noa sø 73ø (73°19.6´n 25°04.6´w). name occasionally used by eha (1953) for a small lake east of noa sø in ymer ø, but first used by andersen (1937) in the form kleine noa see. lille oksedal 72ø-306 (72°00.9´n 23°42.0´w). valley draining from oksehorn into the north side of kolledal. so named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk, because of the numerous musk-ox calves seen here. officially it is considered to be identical to rødedal, although hans kapp evidently regarded lille oksedal as a minor valley on the south side of his rødedal. lille pendulum 74ø-1 (74°40.0´n 18°28.0´w; maps 2, 4). island ne of sabine ø, part of the pendulum øer group. named by karl koldewey’s 1869–70 expedition as kleine pendulum insel (fig. 6), possibly an unfortunate choice of name as the original pendulum experiments were carried out on the present sabine ø. it may corre245 spond to the area designated by william scoresby jr. in 1822 as gale hamke’s land. (pendulum insel, kleine pendulum, lille pen dulum ø, little pendulum island, pendu lumeiland.) lille petermann 73ø-715 (73°04.3´n 28°40.4´w; see also figs 65, 69). pronounced peak 2700 m high on the west side of norden skiöld gletscher, sw of petermann bjerg. named by j.m. wordie’s 1929 expedition as little peter mann, and approved in the 1950s at the suggestion of john haller. lille ravnefjeld 71ø-346 (71°41.1´n 22°44.5´w). mountain 3 km sw of ravnefjeld, wegener halvø. named during the lauge koch expeditions in the 1950s by rudolf trümpy. lille skibssø 76ø-349 (76°46.4´n 18°43.4´w). small lake at dan markshavn, immediately sw of skibssø. the name was sug gested by hans meltofte in 1972, who also noted that the lake was often referred to by the staff at danmarkshavn weather station in 1969–71 as lille vandsø. fischer (1983) notes it was also known as fuglesø. lille snenæs 76ø-63 (76°52.8´n 19°41.1´w; map 4). peninsula east of lumskebugten on the south coast of germania land. so named by the 1906–08 danmark-ekspeditionen, because it was often confused with nearby snenæs (thostrup 2007). it is now a noted haul-out locality for walrus. up to 48 walruses have been recorded here at one time (born et al. 1997). (little snow naze.) lille snenæshytten 76ø (76°52.8´n 19°37.9´w). danish hunting hut at lille snenæs on the south coast of germania land. built by nanok in october 1939. lille stu 73ø (73°26.8´n 27°07.6´w). small norwegian hunting hut at the head of isfjord, on the east side of gerard de geer gletscher, built in march 1940 for arktisk næringsdrift (lille stu = small room). it has also been known as isfjordhytten. lille sødal 74ø-300 (74°19.3´n 20°07.5´w). valley in south wollaston forland where there are many small lakes. the name originated from the wintering parties at kulhus and eskimonæs during the 1931–34 treårsekspeditionen. lille vandsø 76ø (76°46.4´n 18°43.4´w). name reported by hans meltofte as in use by the staff at danmarkshavn weather station in 1969–71 for lille skibssø. lille vandsø 74ø (74°19.4´n 20°11.2´w). hut built by daneborg weather station at a small fresh-water lake in lille sødal, wollaston forland. lillebittesødal 74ø (c. 74°20´n 20°10´w). name used by daneborg weather station personnel for a side valley to lille sødal, wollaston forland (bitte = diminutive; lillebitte = very small; lillebittesødal = very small lake valley). lilledal 72ø-392 (72°02.9´n 23°18.9´w). minor tributary valley to slugtdal, west of antarctic havn. the name was used by hans kapp during lauge koch’s 1957–58 expeditions. lilleelv 72ø-232 (72°40.0´n 22°50.8´w). small river on ne traill ø draining into vega sund. named by desmond t. donovan during lauge koch’s 1949–50 expedition. lillefjord [kangertivatsaakajik] 70ø-212 (70°37.8´n 21°40.7´w; maps 3, 4). fjord on the east coast of southern liverpool land. the name first appears as lille fjord on a map compiled by janus sørensen (sørensen 1928). lillegletscher 71ø (71°58.7´n 26°32.9´w). name occasionally used for a minor glacier between toscano gletscher and sydgletscher, on the south side of northern frederiksdal, nathorst land (zweifel 1958). lillegletscher 75ø-84 (75°59.3´n 22°09.8´w). glacier west of the head of bessel fjord. the name appears to have been suggested by the place name committee in 1935, probably as a replacement for a proposed name they considered unsuitable. limfjordsbakkerne 76ø (78°46.3´n 18°45.1´w). eastern slopes of harefjeldet, near danmark havn. the name was used by friis (1909) in his popular account of the 1906–08 danmarkekspeditionen, because the slopes resembled the locality of the same name in denmark after a heath fire. lindauer hörnli 71ø (71°48.6´n 25°00.5´w; map 5). mountain about 2000 m high on the sw side of roslin gletscher. climbed by karl m. herligkoffer’s 1966 expedition on 21 august, and named after lindau, a town at the east end of bodensee, of which the old town centre dating from the middle ages is built on an island. lindbergh fjelde 69ø-35 (69°07.0´n 30°50.0´w). nunatak area west of christian iv gletscher, northern christian ix land. mapped by lauge koch during flights in 1933 on the 1931–34 treårsekspeditionen, and named lindbergh land after colonel charles lindbergh and his wife, whom koch met on ella ø in august 1933. the lindberghs had flown across the inland ice from the west coast of greenland, and discussed with koch the new land lindbergh had seen. charles augustus lindbergh [1902–74] was best known for the first solo flight across the atlantic ocean in 1927. (lindbergh fjælde; lindbergh nunatakker .) lindbergh gletscher 69ø-45 (69°08.0´n 30°32.0´w). glacier in northern kong christian ix land, named by lawrence wager’s louise boyd landlille cervinfrænkel land jæ tte gle tsc he r fig. 58. looking west over the 4 km wide jættegletscher that separates louise boyd land from frænkel land. the mountain lille cervin in north frænkel land is 1600 m high. the john haller photograph collection, geus archive. 246 1935–36 expedition as lindberghs glacier after nearby lindbergh fjelde. lindeman fjord 74ø-33 (74°40.0´n 20°45.4´w; maps 2, 4). fjord sw of kuhn ø. named by karl koldewey’s 1869–70 expedition as lindeman bai (fig. 6), after moritz karl adolf lindeman [1823– 1908]. he was secretary of the bremen geographical society, the ‘verein für die deutsche nordpolarfahrt’, and editor with g. hartlaub of koldewey’s narrative of the expedition. during the 1931–34 treårsekspeditionen the ‘bay’ was found to extend inland as a 20 km long fjord, and the name amended accordingly (fig. 15). the eastward extension of the fjord between kuhn ø and wolla ston forland is still occasionally re ferred to as lindemans bugt. (lindemann bay, lindemans fjord.) lindeman fjord hytten 74ø (74°38.6´n 20°49.2´w). danish hunting hut on the south side of lindeman fjord, built by nanok in 1931, and rebuilt in 1938. it was burnt down in december 1978. it was also known as fjordhytten. a norwegian hut nearby is known as svendsby. lindemannelv 74ø (74°34.7´n 20°42.4´w). river in lindemans dalen, north of zackenberg forskningsstation. the name is used as a reference locality by visiting scientists. lindemannhytten 74ø (74°45.2´n 20°37.0´w). danish hunting hut on the west side of fligely fjord, north of the mouth of lindeman fjord, built by nanok in 1931. it is also known as sydlige fligely hytten. (lindeman hytten.) lindemans bugt 74ø (74°42.0´n 20°30.0´w). koldewey’s original lindeman bai between kuhn ø and wollaston forland was renamed lindeman fjord when the extent of the ‘bay’ became clear during the 1931–34 treårsekspeditionen. however, vischer (1943) and maync (1947) both used lindemans bugt on their maps for the east extension of lindeman fjord south of kuhn ø, and this unapproved usage has subsequently been perpetuated by its use in a formal stratigraphical division. lindemansdalen 74ø-157 (74°34.7´n 20°42.4´w). valley running from lindeman fjord southwards to young sund. the name first appears in a geological report by frebold (1932). (lindemann dal, lindemanndal.) lindemanspashytten 74ø (74°35.3´n 20°43.6´w). danish hunting hut built by nanok in may 1951 on the west side of the pass at the southern end of lindemansdalen. lindemanssø 74ø (74°30.9´n 20°38.5´w). lake in the sw part of lindemansdalen. the name is used as a reference locality in orni thological reports by visiting scientists to zackenberg forsk nings station. lindhard ø 76ø-118 (76°31.5´n 22°10.0´w; map 4). island south of borgfjord. named by j.p. koch’s 1912–13 expedition as lind hards ø, after jens peter johannes lindhard [1870–1947], doctor on the 1906–08 danmark-ekspeditionen. (lindhards ey.) lindqvist-hytta 72ø (72°53.9´n 24°22.7´w). norwegian hunting hut on the north side of vega sund, se of svedenborg bjerg (nsiu 1932c), built by arktisk næringsdrift in 1929. the name was given for gustav lindqvist, a norwegian hunter who in addition to hunting in east greenland had spent 17 years in spitsbergen. the hut has also been known as nils hermans hytta and nansen-hytten. lindsay nunatak 69ø-40 (69°15.0´n 33°04.0´w). nunatak nw of the prinsen af wales bjerge, northern kong christian ix land, originally named by l.r. wager’s 1935–36 expedition in the form lindsay nunataks to cover a group of three. on the northernmost nunatak l.r. wager had found a broken ice axe left by martin lindsay’s 1934 expedition. wager’s map (wager 1947) shows the location some distance nw of the prinsen af wales bjerge, but a much closer placing is favoured on modern accurate maps (e.g. brooks et al. 1996; nielsen et al. 2001, fig. 3). the locality was visited by a helicopter party in 1995, and more detailed geological studies were made in 2000. martin lindsay [1905–1981] came to fame in the 1930s when he led a series of expeditions to greenland. he later became a conservative member of parliament, and was awarded a baronetcy in 1982. (lindsays nunatak.) lingularyggen 70ø-45 (70°40.8´n 25°18.6´w). minor ridge nw of kap leslie, east milne land, between glaukonitbjerg ansd slottet. it was named during the 1931–34 treårsekspeditionen by her mann aldinger as lingularücken or lingula rücken after the fossil brachiopod lingula. linné gletscher 72ø-243 (72°18.7´n 24°56.2´w; map 5). large glacier in the northern stauning alper, named by erdhardt fränkl during lauge koch’s 1950–51 expeditions after carl von linné [1707–1778]. linné (or carol linnaeus) was a noted swedish bota nist and explorer, who framed the principles for defining genera and species. lisbet ø 71ø (71°15.7´n 24°55.8´w). small island south of sydkap, one of the present immikeertivaqqat. it was named in 1937 by aage gilberg after his girlfriend, later his wife, in the course of archaeological excavations (gilberg 1987). glob (1946) referred to the island as ruin ø. gilberg also built a cairn on the island referred to as lisbets varde. lisbeth thora gilberg [1917–1992] set a record in 1939 for the farthest north then reached by a european woman with a dog sledge (79°n), and made anthropological studies of polar inuit. listerudodden 73ø (73°27.2´n 21°18.3´w). small peninsula at the mouth of annielven on the south coast of hold with hope. so named on the 1932a nsiu map, after the telegraphist, named listerud, who had manned the first myggbukta radio station in 1921–22. he was lost with the other members of the expedition when the anni 1 was crushed in the pack ice in 1922. listhaugøya 72ø (72°45.6´n 22°48.0´w). small island in vega sund, part of the present scott keltie øer group. so named on the nsiu maps of lacmann (1937), after the norwegian hunter johan listhaug [b. 1910], who wintered in east greenland from 1933 to 1935. little chocolate mountain 73ø (73°21.0´n 25°07.9´w). prominent ridge north of noa sø, west ymer ø, the present rosinante. the name was given by a.b. cleaves and e.f. fox in the course of geological work during john k. howard’s 1933 expedition, for the chocolate-brown colour of the rocks. little cumbrae 71ø (71°56.4´n 25°10.6´w; map 5). small glacier, an upper branch of cantabræ, stauning alper. so named by the 1998 scottish mountaineering club expedition. liverpool land 70ø-149 71ø-121 (71°00.0´n 22°00.0´w; maps 3, 4; see also fig. 72). mountainous land area bounded to the west by hurry inlet, klitdal and carlsberg fjord, and extending from latitude 70°27´n to 71°31´n. william scoresby jr. in 1822 originally gave the name the liverpool coast to the south and east sides of the tract of land now known as liverpool land, because its headlands and islands had been chiefly named after liverpool friends. nordenskjöld (1907) considered the name inappropriate and changed it to liverpool land. (liverpool coast, liverpool kyst, liverpool kusten, terre de liverpool, liverpoolland, côte de liver pool, liverpool kuste, liverpoolküste.) lizard peak 73ø (73°34.3´n 25°54.9´w). subsidiary peak on the south side of grejsdalen, andrée land, on which a series of rock climbs were made. climbed by the 2007 army boreal zenith expedition. lloyds point 70ø (c. 70°38´n 22°36´w). a prominence in hurry inlet, it was named by william scoresby jr. in 1822 after the captain of the trafalgar, who had made useful investigations in the area. it was probably a point on the west side of hurry inlet, but was not depicted on scoresby’s (1823) map, and the name has not been approved. loch fyne 73ø-16 74ø-268a (73°48.0´n 21°48.5´w; maps 2, 4). n–s-trending fjord between hold with hope and hudson land. it was explored by douglas clavering in 1823 and named loch fine after the fjord of the same name in scotland. maps of scotland 247 used the form ‘loch fine’ until at least the middle of the 19th century, whereas the modern spelling is loch fyne. loch fine was used on norwegian maps of east greenland in the 1930s. the change in spelling seems to date from the maps of j.m. wordie’s 1926 and 1929 expeditions. (loch-fine, loch fine fjord.) loch fyne station 73ø (73°40.7´n 21°51.4´w). danish hunting station built by nanok in 1945 on the west side of inner loch fyne. it was manned from 1945 to 1951, and subsequently maintained by sirius. it is considered to be one of the best stations on the coast, which has given rise to the alternative and flattering name of kystens perle. (loch fyne, loch fyne-hytten.) loddevig 76ø-93 (76°43.0´n 18°34.7´w). small bay south of danmark havn, so named by the 1906–08 danmark-ekspedi tionen because detailed soundings were made here (j. løve, personal communication 2009; lod = a sounding weight). lodin elv 71ø-195 (71°22.8´n 24°00.0´w; map 4). river in jameson land draining sw to hall bredning. the name was one of a group of names given by the place name committee in 1939 to replace names proposed by hans stauber. it was given for lodin, who brought home the body of finn fegin from greenland about 1028 after he was lost with his ship. (lodins elv.) lodlineklippe 76ø-343 (76°22.0´n 23°55.4´w; map 4). near vertical cliff on the south side of budolfi isstrøm, south dronning louise land. named by the 1952–54 british north greenland expedition as lodlinieklippe because it was so nearly vertical that a plumb line (= lodline) could be dropped from top to bottom. lollandselv 70ø-92 (70°53.5´n 24°00.0´w; map 4). river in jameson land flowing west to hall bredning. so named by laurits bruhn during the 1931–34 treårsekspeditionen after the island of lolland, denmark. lommensø hytten 71ø (c. 71°48´n 24°20´w). hut built by nordisk mineselskab not far from the lomsøen airstrip, where pingo dal meets schuchert dal. it is also known as pingo dal hytten. the hut was removed in 1990 by a nordisk mineselskab clear-up team. lomsø 76ø-237 (76°48.6´n 19°10.5´w). small lake on winge kyst. so named by the 1906–08 danmark-ekspeditionen after the red-throated diver (= rødstrubet lom), a common breeding bird in the region. (lomsøen.) lomsø 74ø (74°27.5´n 20°33.3´w). minor lake south of zacken berg forskningsstation, close to young sund. the name is used as a reference locality by visiting scientists (meltofte & thing 1996). (loon lake). lomsøen 71ø-292 (71°48´n 24°14´w; map 4). lake in the pass west of the head of pingo dal, which became known during lauge koch’s 1953–54 expeditions under the name lummen sø (bearth 1959). it was named after the red-throated diver which breeds commonly in small lakes. a small airstrip was built adjacent to the lake in 1957 by nordisk mineselskab after a 150 km spring journey by bulldozer and sledge with 70 tons of equipment. lonetoppen 71ø-410 (71°55.0´n 23°14.1´w). mountain 1131 m high in claudius clavius bjerge, north of ørsted dal. named by katharina perch-nielsen during the 1967–72 ggu scoresby sund expeditions after lone malmros who worked in the region in 1969 (see also malmros klint). loon lake 71ø (71°21.3´n 24°48.9´w). name used by hall (1963, 1966) for a small lake at the east end of holger danske briller where the red-throated diver (loon) was observed to nest by the 1962 oxford university expedition. loon lake 72ø (72°53.3´n 22°08.6w). lake on geographical society ø where samples were collected for radiocarbon age determinations (cremer et al. 2008). loppa 72ø (72°39.4´n 22°33.6´w). very small island in vega sund, west of nordenskiöld ø. so named on the nsiu maps of lacmann (1937), for its diminutive size (loppa = flea). louise boyd land 73ø-590 (73°33.0´n 27°54.0´w; maps 2–4; fig. 58). land area between gerard de geer gletscher and jætte gletscher. mapped by lauge koch during flights in 1932 on the 1931–34 treårsekspeditionen, and named miss boyd land after louise arner boyd [1887–1972]. an american polar explorer, she led seven expeditions to the arctic, four of which were to east greenland, and in 1931 was the first to penetrate to the head of isfjord. louise boyd was especially noted for her use of photography, and photogrammetric survey techniques (boyd 1935, 1948). odell (1943) records the ascent of several peaks in louise boyd land during louise boyd’s 1933 expedition. (louise a. boyd land.) louise elv 74ø-117 (74°24.1´n 21°21.8´w). river on nw clavering ø draining into tyrolerfjord, named by lauge koch’s 1929–30 expeditions in the form louise river. girl’s name. a norwegian hunting hut on the west side of louise elv built in 1927 by the foldvik expedition has sometimes been referred to as louise elv hytten, but is more commonly known as bakkehaug. louise gletscher 73ø-609 (73°32.0´n 27°32.0´w). glacier in se louise boyd land, named during louise boyd’s 1933 expedition as louise glacier after the expedition leader (odell 1937a). see also louise boyd land. it was one of the glaciers studied in detail by the expedition. (louises glacier.) luciadal 73ø-403 (73°22.7´n 25°49.5´w). valley in southern andrée land, draining via benjamin dal into eleonore bugt. named during lauge koch’s 1948–50 expeditions by erdhardt fränkl after the popular song ‘santa lucia’. it is said he wanted to give a nice name to a very pleasant valley. luciagletscher 73ø-402 (73°27.0´n 26°00.0´w; map 4). glacier in south andrée land draining via luciadal and benjamin dal to eleonore bugt. named during lauge koch’s 1948–50 expeditions by erdhardt fränkl after luciadal. (lucia gletscher.) ludlams hule 74ø (c. 74°27´n 20°15´w). cave on the east side of brachiopoddal, west wollaston forland. the name was used by rosenkrantz (1932) in his report on geological work during lauge koch’s 1929 expedition. it was named after the 19th century opera ‘ludlams hule’ by adam oehlenschläger. lugano bjerg 72ø-418 (72°48.0´n 27°27.1´w; map 4). mountain in north gletscherland, named during the 1931–34 treårseks peditionen by eugène wegmann as monte lugano, after the swiss town of lugano. it was climbed by eugène wegmann and augusto gansser on 11 august 1934. gansser was from lugano, and is said to have married a girl from one of the best lugano families (fritz schwarzenbach, personal communication 1996). c. mountain and scoop mountain have also been used. lugeon bjerg 72ø-115 (72°38.1´n 25°23.1´w). snow-capped mountain on the west side of polhemdal in south lyell land. so named by eugène wegmann during the 1931–34 treårseks peditionen in the form mont lugeon, after maurice lugeon [1870–1953], a french stratigrapher and structural geologist. for many years he was professor at the university of lausanne, and noted especially for his work on alpine tectonics. (lugeons bjerg.) lumskebugten 71ø (71°55.7´n 28°27.4´w). name used by helge g. backlund (in: koch 1955) for the iceberg-filled inner part of nordvestfjord in front of daugaard-jensen gletscher. probably named for the near-fatal accident to backlund’s party caused when the front of nearby daugaard-jensen gletscher collapsed. (lumsk = treacherous). lumskebugten 72ø-79b (72°53.7´n 25°42.1´w). bay on the se coast of suess land at the mouth of murgangsdal, named by j.m. wordie in 1929 as deceit bugt for its misleading appearance. the flat valley at its head at first sight suggests the bay extends much farther north. mineralbugt has also been used. (deceitbucht.) lumskebugten 76ø-62 (76°55.0´n 19°53.0´w). bay on the south coast of germania land. so named by the 1906–08 danmarkekspeditionen by christian b. thostrup after the bar or cafe just outside the gates of the harbour authority in copenhagen, now the noted restaurant at the same location. thostrup (2007) records 248 that like the bay the cafe had the tendency to attract unwary passers-by. (wily bay.) lumskebugthytten 72ø (72°53.8´n 25°43.9´w). norwegian hut on the west side of the floodplain at the head of lumskebugten. it was built between 1934 and 1938 by arktisk næringsdrift, and was originally known as sunnmøresheimen and later as mineralbukta. lunckefjellet 73ø (73°17.0´n 23°37.0´w). mountain ridge north of dusén fjord, including the present udkiggen. so named on an nsiu map (1932a) after bernhard luncke [1894–1963], a norwegian topographer, and a pioneer and expert in aerial photogrammetry. he took part in 18 expeditions to the polar regions, often as leader, including the nsiu expeditions to east greenland from 1929 to 1933. (mt. luncke, mt. lunke.) lunedal 72ø-516 (72°33.6´n 24°00.3´w). valley on sw traill ø draining south to holm bugt. so named by geoffrey halliday following botanical work during the 1961 leicester university and 1971 northern universities expeditions. origin of name uncertain. lurcher’s crag 77ø (77°13.0´n 24°23.9´w). name used informally by a surveying party of the 1952–54 british north greenland expedition for krebs bjerg, dronning louise land. tripods were twice blown over here while surveying (banks 1957). luxembourg spids 73ø (73°09.0´n 28°30.5´w). name given to the 2517 m high sw peak of trappebjerg in western frænkel land by a dundee university expedition. the peak was climbed in august 1985. lycett bjerg 72ø-237 (72°21.6´n 22°55.2´w). mountain on se traill ø, north of bjørnedal. named by desmond t. donovan during lauge koch’s 1949–50 expeditions after john lycett, a 19th century palaeontologist who worked on fossils of the same age as those that occur in the mountain. lyell land 72ø-39 (72°36.0´n 25°34.0´w; maps 3, 4; fig. 29). land area bounded by kempe fjord, kong oscar fjord and forsblad fjord. named by a.g. nathorst’s 1899 expedition after charles lyell [1797–1875], one of the most influential of british geologists, especially noted for his ‘principles of geology’, published in 1830 and running to 12 editions. (lyells land, lyell-land.) lygnaelv 73ø-148 (73°37.2´n 20°37.8´w). river in se hold with hope, flowing across østersletten. named on an nsiu map (1932a) as lygna, possibly after a river of the same name in the oppland area of norway. lyngedalen 72ø (72°51.6´n 22°48.1´w). valley on central geograph ical society ø draining south into vega sund, equivalent to the present lysdal. so named on nsiu maps of lacmann (1937) after bernt arne lynge [1884–1942], a norwegian botanist who was professor of botany at oslo university. he took part in several arctic expeditions including nsiu expeditions to east greenland. lynn ø 80ø-4 (80°07.8´n 19°12.8´w; maps 1, 4; fig. 24). island bounded by hekla sund and dijmphna sund. so named by the 1906–08 danmark-ekspeditionen. christian b. thostrup re cords that it was named after a british shipping company at bridgeness (thostrup 2007). lysdal 72ø-174 (72°51.6´n 22°48.1´w; map 4). valley on central geographical society ø. the name was one of a group of names given by the place name committee in 1939 to replace proposals by hans stauber. it is also a danish place name. lyngedalen has also been used. lysevig 76ø (76°55.1´n 21°00.0´w). name used in charles poulsen’s diaries of the 1906–08 danmark-ekspeditionen (poulsen 1991) for the side branch of mørkefjord more usually known as puster vig. this short fjord or bay has a lighter aspect than the steep-sided mørkefjord (mørk = dark, lys = light). (lysefjord.) lysevig huset 76ø (76°55.3´n 21°01.6´w). name used in charles poulsen’s diaries of the 1906–08 danmark-ekspeditionen for peter freuchen’s meteorological station in pustervig (also known as lysevig and pustervig). lysippsdal 73ø-699 (73°17.4´n 26°50.0´w). valley on the ne side of frænkel land draining into isfjord. so named by john haller following explorations during lauge koch’s 1949–51 expeditions, because the north wall at the entrance to the valley is said to resemble a statue by the greek sculptor lysippus. lystergletscher 73ø-610 (73°13.8´n 27°43.4´w). glacier in west frænkel land, formed by the merging of three glaciers of about the same size. named by louise boyd in 1933 as trident glacier (lyster = trident). lächen 72ø-448 (72°57.8´n 26°02.1´w). spectacular waterfall and gorge on the south side of inner murgangsdal, suess land. the name was used by eugène wegmann during the 1931–34 treårs eks peditionen, and is a swiss dialect word for a flood or lake. the periodic drainage of murgangssø is through this gorge. lægervallen 79ø-14 (79°14.4´n 18°59.1´w; maps 1, 4). flat cape on east lambert land, north of brønlunds grav. named by the 1938– 39 mørkefjord expedition. it is a nautical expression for a sandy beach on the lea-side. længselsbjerg 73ø-384 (73°39.3´n 25°42.9´w). mountain in andrée land north of grejsdalen, named by erdhardt fränkl during lauge koch’s 1948–50 expeditions. længselsklippen 72ø-170 (72°09.3´n 22°32.4´w). cliff on south traill ø, se of drømmebugten and sw of purpurfjeld. named during lauge koch’s 1936–38 expeditions by hans p. schaub (længsel = longing, yearning). (heimwehfluh.) læsø 72ø-101 (72°35.6´n 22°20.2´w; map 4). island off ne traill ø at the mouth of vega sund. named by ove simonsen during the 1931–34 treårsekspeditionen after the danish island of the same name in the kattegat, se of frederikshavn. løberen 71ø-418 (71°38.5´n 25°30.1´w; maps 4, 5). surging glacier in the south stauning alper, which advanced 7.5 km between 1950 and 1967, when it was observed to have reached nord vestfjord. named by johan d. friderichsen during the 1967–72 ggu scoresby sund expeditions (løber = runner). neptune glacier has also been used. lökvik-hogda 73ø (73°30.8´n 21°35.9´w). hill 252 m high north of myggbukta. so named on an nsiu map (1932a). perhaps named after the norwegian town of løkvik. (lökvik hill.) lögtoppene 74ø (74°22.0´n 19°52.1´w). n–s-trending ridge be tween grænsedalen and blæsedalen in south wollaston forland. the name was used by wolf maync and andreas vischer during their work on lauge koch’s 1936–38 expeditions, and is found on vischer’s (1943) maps. the mountain summits may be likened to the shapes of onions (= løg). løvebastionen 73ø-543 (73°01.0´n 28°33.0´w; see also fig. 69). prominent crag 2500 m high on the south side of nordenskiöld gletscher, named by j.m. wordie’s 1929 expedition as lion bastion for its appearance (løve = lion). (lövebastionen.) løvehovedet 73ø-648 (73°50.1´n 25°22.1´w). mountain 902 m high in west strindberg land, on the ne side of geologfjord. so named during the 1931–34 treårsekspeditionen by th. johansen because of a resemblance to the løvehovederne in north bornholm (løvehovedet = the lion’s head). løyningdalen 72ø (72°56.1´n 23°56.7´w). valley on west geograph ical society ø draining south into vega sund. used on the nsiu maps of lacmann (1937), and named after paul løyning [1895– 1960], a norwegian zoologist who became curator at the zoo logical museum in oslo in 1926, and took part in nsiu expeditions to east greenland from 1930 to 1932. (löyningdalen.) låg-øyra, laag-øyra – see lavøira. låghumpen 73ø (73°31.2´n 21°04.8´w). hill 315 m high in south hold with hope. the name appears on an nsiu map (1932a; fig. 13, and was probably given for its relatively low, humpy nature. m mackenzie bugt 73ø-9 (73°27.0´n 21°30.0´w; maps 2–4). bay 249 south of hold with hope. william scoresby had seen an opening of the land at a great distance in 1822, and named it mackenzie’s inlet in compliment to sir george steuart mackenzie [1780–1848]. a mineralogist, noted for his proof of the identity of diamond with carbon, mackenzie was, like scoresby, a pupil and friend of robert jameson. karl koldewey in 1869 observed the supposed inlet to be a bay. norwegian hunters used myggbukta for the same feature in the 1920s and 1930s, but this name was later restricted to the norwegian radio station in the bay. (mackenzies inlet, mackenzie bay, mackenzie einbucht, mackenzie bugten, mackenziebukten, mackenzie-bai.) mckenzie glacier 71ø (71°59.7´n 25°17.0´w; map 5). minor branch of essemmceebrae on the south side of sefström gletscher, stauning alper. so named by the 1998 scottish mountaineering club expedition. mackenzie river 73ø (73°30.0´n 21°44.8´w). river draining through badlanddal into mackenzie bugt. the name was used by goodhart & wright (1958). mackenzie valley 73ø (73°34.0´n 21°48.0´w). valley north of mackenzie bugt, the present badlanddal. the name was used in reports of louise boyd’s 1933 expedition (boyd 1935). macknight bjerg 71ø-28 (71°23.3´n 22°31.7´w; map 4). mountain 540 m high on the west side of carlsberg fjord. named by william scoresby jr. in 1822 as cape macknight, probably after thomas macknight [1763–1836], minister of the old church edinburgh from 1810. like many of scoresby’s capes it was later shown to be a mountain. (cape machnight.) maclear 71ø (71°39.1´n 25°13.0´w; map 5). mountain about 1900 m high on the north side of mercurius gletscher, south stauning alper. first climbed by james clarkson’s 1961 expedition. madum sø 73ø-646 (73°44.1´n 27°23.9´w). lake in west andrée land at the north margin of gerard de geer gletscher. so named by ove simonsen during the 1931–34 treårsekspeditionen because, like the danish lake of the same name in ne jylland, it has no apparent exit. magdalenasø 73ø-355 (73°52.2´n 25°20.6´w; map 4). lake at an altitude of 585 m in west strindberg land, named by hans r. katz during lauge koch’s 1948–49 expeditions. magdalene spids 72ø-499 (72°07.8´n 25°13.5´w; map 4). highest peak west of vertebrae, on the north side of gully gletscher, stau ning alper. climbed by the 1963 cambridge university expedition on 18 august, and named after magdalene college, cambridge, founded in 1542. (magdalene.) magga dan gletscher 69ø-81 (69°55.0´n 27°05.0´w; maps 3, 4; fig. 22). larger of two large glaciers which flow northwards into gåsefjord. named by w. stuart watt during the 1967–72 ggu scoresby sund expeditions after the expedition ship magga dan used in 1969. the magga dan, built in 1956 for the j. lauritzen shipping company was a noted polar expedition ship. magga ø 72ø-330 (72°43.4´n 22°51.8´w). small island adjacent to kista ø in vega sund. the name was proposed by søkortarkivet in 1956–57 following surveying of the channel through vega sund as an alternative approach for ships en route to mestersvig, and given for the magga dan. see also magga dan gletscher. magnetikerbjerg 74ø-108 (74°10.4´n 20°13.8´w). mountain on east clavering ø, named by lauge koch’s 1929–30 expeditions in the form mt. magnetiker. probably named after max grotewahl, a member of the 1930 expedition who carried out geomagnetic observations in the region. magog 71ø (71°55.7´n 25°07.2´w; map 5). mountain with twin summits at the head of the easternmost branch of canta bræ. it was reported as shaped like the head and beak of a bird. the 1963 cambridge university expedition climbed the east spire on 8 august, and in some of their reports refer to it as gog magog. see also magog below. magog 73ø-535 (73°15.8´n 28°22.2´w; map 4; see also fig. 65). mountain 2400 m high in west frænkel land. it was first climbed by w. huber and hans r. katz on 25 august 1948. the name had been given by j.m. wordie’s 1929 expedition together with its slightly higher neighbour gog, for the gogmagog hills near cam bridge. see also gog. main glacier 71ø (71°46.5´n 25°13.4´w; map 4). name used by john hunt’s 1960 expedition for the upper section of the present bjørnbo gletscher, south stauning alper. the name is used in mountaineering literature for the main branch of bjørnbo glet scher nw of concordia. majdal 72ø-382 (72°01.2´n 23°22.5´w). valley in north scoresby land west of antarctic havn, draining south into kolledalen. named by hans kapp during lauge koch’s 1957–58 expeditions. majhytten 76ø-208 (76°17.1´n 21°07.1´w). danish hunting hut on the north side of the mouth of syttendemajfjorden, built by nanok in september 1938. now a ruin. (17. maj hytten, syttende maj hyt ten.) major paars dal 71ø-190 (71°32.8´n 24°11.0´w). valley in western jameson land draining sw into schuchert dal. the name was one of a group given by the place name committee in 1939 to replace proposals by hans stauber. it was given for major claus enevold paars, a dane whom frederik iv sent to greenland as governor, and whose name is best known for a failed attempt to cross the inland ice on horseback. majorpasset 72ø-362 (72°06.8´n 24°54.8´w; map 5). pass 2150 m high between bersærkerbræ and gulley gletscher, the key pass to the traverse of the central stauning alper. it is better known in mountaineering literature as col major, the original name proposed by malcolm slesser in 1958 who made the first crossing. majskær 70ø-228 (c. 70°45´n 21°26´w). group of skerries off the coast of south liverpool land. the name first appeared on a map compiled by janus sørensen in the form majskærene (sørensen 1928). malia havn 72ø-335 (72°41.7´n 22°37.9´w). small harbour on south geographical society ø, adjacent to kap hovgaard. the name was proposed by søkortarkivet in 1956–57 following surveying of the channel through vega sund as an alternative approach for ships en route to mestersvig. jomfrupollen was used for the same feature by lacmann (1937). mallemukfjeld 80ø-6 (80°11.8´n 16°37.9´w; map 4). cliff in se holm land, named by the 1906–08 danmark-ekspeditionen as mallemukfjeldet because of the large colony of fulmars observed here in april 1907. eigil nielsen (1941) noted that the 1906–08 danmark-ekspeditionen were inconsistent in their usage, sometimes applying the name to the present depotfjeld, and more precisely defined the name to apply to the most precipitous of the cliffs. (mallemukfjæld, mallemuk hill, mallemukfjället.) mallemukgletscher 80ø-34 (80°13.0´n 16°38.2´w). glacier on the north side of mallemukfjeldet in se holm land. named during the 1909–12 alabama expedition after nearby mallemukfjeld. mallemukken 80ø (80°08.5´n 22°30.5´w). sirius hut on the south shore of centrumsø, built by slædepatruljen daneborg on 8 august 1952. it was in regular use until may 1979, when it was replaced by the modern sirius hut at the west end of centrumsø. (mallemuk hytten.) malmbjerg 71ø-260 (71°57.4´n 24°16.7´w). mountain between schuchert gletscher and arcturus gletscher, named by peter bearth and eduard wenk during lauge koch’s 1953–54 expeditions. the name is usually applied to the conspicuous rust-red, black and yellow colours on the sw flank of the mountain due to mineralisation (molybdenum, wolfram, galena, zinc and pyrite). the first drilling in 1958 was followed up by extensive drilling in 1961–62. in all, 147 boreholes totalling 22 877 m were drilled and 1329 m of adits excavated. an ore deposit of 150 million tons with a grade of 0.23% mos2 and 0.02% wo3 was proven (harpøth et al. 1986). the remains of the drilling camp stood until the late 1980s 250 on the moraine, but have now been demolished. arktisk mine kompagni held a concession to mine and ship molybdenum from 1961 to 1984, but due to the low grade the deposit was not exploited. swedish geophysical companies involved in the evaluation work usually used the form erzberg. the dramatic price increases of metals led to initiation of a new phase of evaluation in 2005, that was put ‘on hold’ in 2008 with the world-wide financial crisis and a slump in metal prices. malmquist plateau 74ø-232 (74°09.9´n 20°41.0´w). small plateau on se clavering ø, west of moskusokseelv. named by lauge koch after david malmquist [b. 1904] who undertook prospecting in the region during the 1931–34 treårsekspeditionen. (malmquists pla teau.) malmros klint 71ø-413 (71°42.5´n 23°04.7´w). cliff on the nw side of fleming fjord. named by katherina perch-nielsen during the 1967–72 ggu scoresby sund expeditions after lone malmros [d. 1969], a geologist who worked in the area in 1969, and died in a car-accident in denmark shortly after returning home. manby halvø [pukkitsivakajik] 69ø-5 (69°49.0´n 23°04.0´w; maps 3, 4). peninsula on the northern part of the blosseville kyst. william scoresby jr. named manby island in 1822 after george william manby [1765–1854], in gratitude of his exertions and success in the rescue of ship-wrecked mariners. manby had developed an early form of breeches-buoy, which up to 1823 had saved 229 lives. in 1821 manby accompanied scoresby on a whaling voy age. scoresby’s island was subsequently shown to be a peninsula (amdrup 1902b). (manby ö.) manley bjerg 74ø-143 (74°15.0´n 22°32.6´w). mountain 960 m high south of grantafjord. lauge koch’s 1929–30 expeditions originally gave the name manley land to the area west of copeland fjord (fig. 15) corresponding to the present blosseville bjerg, courtauld bjerg and manley bjerg, because the area was first mapped by gordon manley during j.m. wordie’s 1926 cambridge expedition. backlund (1932) used the name for the peninsula of which blosseville bjerg is the highest point. gordon manley [1902–80], a geographer who made notable con tri butions to meteorology and climatology, was professor at bedford college from 1948 to 1964 and later professsor at the university of lancaster. manley land 74ø (74°15.7´n 22°11.1´w). name used on lacmann’s (1937) maps for the present blosseville bjerg west of clavering ø. see also manley bjerg. manniche sø 76ø-348 (76°12.5´n 21°17.0´w; map 4). lake in northern ad. s. jensen land. named during lauge koch’s 1956–58 expeditions by john haller, after a member of the 1906–08 dan mark-ekspeditionen. arne ludvig valdemar manniche [1867– 1957] was ornithologist on the expedition, and subsequently wrote several handbooks on denmark’s birds. marabugt 72ø-274 (72°50.4´n 24°53.2´w). bay on east ella ø. named by john w. cowie during work carried out from 1949 to 1954 on lauge koch’s geological expeditions, possibly after the wife of peter adams. marcia bjerg 73ø-691 (73°23.5´n 26°31.1´w; map 4). mountain c. 1460 m high in sw andrée land, between rendal and jomsborg dal. named by john haller following explorations during lauge koch’s 1949–51 expeditions, after the mountain ‘la marcia’ in switzerland, which has a similar shape and similar rocks, and is also situated at a fork in a valley. margaret lambert sø 71ø (71°30.5´n 25°20.7´w). name used for the eastern of the two lakes of holger danske briller on 1952 wac maps. margaretasø 73ø-690 (73°25.5´n 26°39.0´w). lake in rendal, sw andrée land. so named by john haller following explorations during lauge koch’s 1949–51 expeditions, because the reflection of margaretatop could be seen in it. both features were named after margareta hediger. margaretatop 73ø-694 (73°23.1´n 26°13.3´w). mountain about 2360 m high in southern andrée land. named by john haller following explorations during lauge koch’s 1949–51 expeditions, after margareta hediger. (margarita spids.) margarinecentralen 76ø (76°56.5´n 18°10.8´w). hunting hut at kap steensby on the east coast of germania land, built in august 1938 by the norsk–franske polarekspedisjon. the expedition had been given a large quantity of margerine, mainly used as dog food, and the hut was built with the empty boxes. the hut has also been known as kap steensby hytten and resoluthytten. (centralen.) margerie dal 73ø-626 (73°09.6´n 25°55.5´w). valley on sw ymer ø, named during the 1931–34 treårsekspeditionen by eugène wegmann as margerie valley after emmanuel de margerie [1862– 1953], a noted french geologist and geographer. he was an honor ary professor at the university of strasbourg, librarian to the société géologique de france and a foreign member of the royal society. (margeries dal.) margretabjerg 71ø (71°58.6´n 24°51.0´w; map 5). mountain 2430 m high on the west side of upper storgletscher with an m-shape as seen from the sw, central stauning alper. climbed and named by the 2007 smc east greenland expedition; the name was given for margaret n. litterick [1927—2005]. margrethedal 73ø-85 (73°19.8´n 22°34.1´w). valley on south gauss halvø, named during lauge koch’s 1929–30 expeditions in the form margrethe valley, after the wife of richard bøgvad. bøgvad and arne noe-nygaard worked here in august 1930. nor we gian maps use the name smedal or smedal valley. margrethedalhytten – see smedal. maria ø 72ø-47 (72°57.3´n 24°53.7´w; map 4). island north of ella ø at the mouth of kempe fjord. named by a.g. nathorst in 1899 after his daughter ella maria charlotte [b.1881], in the form marias ö (fig. 8). see also ruth ø and ella ø. (maria island, maria øya, maria-öya, marie island.) mariager fjord 70ø-253 71ø-127 (70°59.1´n 21°52.5´w; map 4). fjord on the east coast of liverpool land. named during the 1931– 34 treårsekspeditionen by laurits bruhn after the fjord of the same name on the east coast of jylland, denmark. marianne nunatakker 74ø-141 (74°34.8´n 23°37.7´w). group of nunataks in wordie gletscher, named by lauge koch’s 1929–30 expeditions as the marianne nunataks. they were visited by a geological party in 1932. girl’s name. mariannes see 74ø (74°35.´n 23°26.9´w). lake east of marianne nunatakker on the ne side of wordie gletscher. the name was used by mittelholzer (1941), and also appears on ams maps. marie-theresia bjerg 72ø-338 (72°27.2´n 22°10.2´w). mountain on se traill ø, so named during lauge koch’s 1956–58 expeditions by h.p. heres after marie-thérese of austria [1638–83], consort of louis xiv of france. maristua 72ø (72°53.6´n 24°47.3´w). norwegian hunting hut on ne ella ø, 3 km south of kap elisabeth, built by arktisk nærings drift in 1930. the hut has also been known as camp lindquist. märjelen see – see after mæchel-stua (‘ä’ is treated as ‘æ’ in danish). markusdal 71ø-321 (71°36.5´n 24°52.8´w). minor valley draining into gurreholm dal, west of schuchert dal. named by enrico kempter during lauge koch’s 1956–58 expeditions, after his coworker markus aellen. marmorbjerg 72ø-287 (72°34.6´n 27°28.6´w; map 4). mountain in west gletscherland, traversed by several thick marble bands. named by john haller following explorations during lauge koch’s 1952–53 expeditions (marmor = marble). marmorknold 70ø-444 (70°15.2´n 29°26.2´w). north point of an 890 m high nunatak on the se side of vestfjord gletscher. so named by w.e.a. phillips during the 1967–72 ggu scoresby sund expeditions because it was formed of yellow-white marble. marmorvigen 80ø-59 (80°05.4´n 20°05.2´w; maps 1, 4; fig. 24). bay on the west side of hekla sund, kronprins christian land. so named by elmar drastrup’s 1938–39 expedition because yellow 251 marble (= marmor) crops out at the head of the bay. maroussia 76ø-37 (76°39.5´n 18°30.6´w). small island east of lille koldewey where the duke of orléans landed on 26 july 1905. he named it after his yacht, the maroussia, used previously on voyages to svalbard in 1896 and 1904. (îlot maroussia, maroussia ø, maroussia island, maroushia.) marrakajik [schuchert flod] 71ø-59 (71°17.3´n 24°36.9´w). extensive muddy delta area at the head of nordostbugt, the lower part of the schuchert flod braided river system. recorded by the 1955 geodætisk institut name registration, the name translates as ‘the small clay’. this greenlandic name has also been used for adjacent nordostbugt. (maqqakajik.) mars glacier 71ø (71°13.3´n 26°17.1´w). glacier on the north side of edward bailey gletscher, renland. named by the 2007 west lancashire mountaineering group expedition. mars gletscher 71ø-335 (71°45.2´n 25°00.7´w; map 5). glacier on the north side of bjørnbo gletscher, south stauning alper. named mars glacier by john hunt’s 1960 expedition, after mars, the fourth major planet from the sun. mars tooth 70ø (70°55.0´n 25°50.3´w). tooth-like summit about 1500 m high on northern milne land. it was climbed by the 1989 greenland milne land expedition. marstranderfjellet 74ø (74°13.3´n 21°18.3´w). mountain 1162 m high on sw clavering ø, the present vestmar bjerg. the name is used on the nsiu maps of lacmann (1937), and was given for fredrick marstrander [b. 1915], who took part in the 1932 nsiu expedition to east greenland. martaajik 70ø (c. 70°32´n 23°38´w). name used by tuborg & san dell (1999) for an inuit ruin site on the coast of southern jameson land, at the western mouth of the river draining flakkerhuk. martin karlsen bugt 71ø-385 (71°30.0´n 27°11.5´w). prominent bay on the south side of central nordvestfjord. named during the 1967–72 ggu scoresby sund expeditions after the expedition ship used in 1968, the martin karlsen, formerly the kista dan. the martin karlsen was named after the noted norwegian shipping company of the same name. see also martin karlsensundet and kista dan gletscher. martin karlsen dal 71ø-386 (71°28.0´n 27°31.0´w; map 4). valley in th. sørensen land, draining into martin karlsen bugt. named during the 1967–72 ggu scoresby sund expeditions. see also martin karlsen bugt. martin karlsensundet 72ø-n311 (72°42.1´n 22°49.1´w). sound between thora ø and silja ø in vega sund. so named on the nsiu maps of lacmann (1937) after martin karlsen [b. 1892], noted norwegian ship-owner whose main activities were sealing in arctic waters. martin knudsen nunatakker 73ø-589 (73°18.0´n 29°04.0´w; map 4). nunatak area west of victor madsen gletscher. mapped by lauge koch during flights in 1932 on the 1931–34 treårseks peditionen, and named after martin knudsen [1871–1949]. knudsen was professor at the university of copenhagen from 1912 to 1941, leader of danske hydrografiske undersøgelse (danish hydrographical survey) from 1902, and was on the committee of the 1931–34 treårsekspeditionen. the original usage was broader than the present, and included nils holgersen nunatakker to the west. (martin knudsens nunataks.) maryhuset 74ø (74°09.9´n 20°11.7´w). norwegian hunting station at kap mary on se clavering ø, built in august 1909 by vebjørn landmark, and subsequently used by the 1927–29 hird expedition and arktisk næringsdrift. dangerous ice conditions that caused the death of three hunters led to a decision to demolish the hut in 1947, the materials being used to build dahls skær hytten. østgrønlandske fangstkompagni built a house back-to-back with the norwegian house in 1921, but this was dismantled in 1923 (see also christianshavn). (mary-huset, kapp mary, kap mary huset.) masclet bay 71ø (71°05.4´n 21°54.6´w). fjord in liverpool land, now known as storefjord. the name masclet bay was given by william scoresby jr. in 1822 to what appeared to be a small bay or inlet, and was named after the late french consul at liverpool, chevalier masclet. the name is not given on scoresby’s chart, though it can be clearly identified from the description in the text and the appendix. both capes guarding masclet bay have retained scoresby’s original names. (masclet bucht.) matterhorn 75ø-16 (75°25.1´n 20°53.6´w; map 4; fig. 59). moun tain 1624 m high in the southern barth bjerge, north of ardencaple fjord. named by karl koldewey’s 1869–70 expedition after the mountain of the same name in switzerland. several names in the region were derived from swiss or austrian mountains because of their alpine aspect. matterhorn was climbed in 1952 by members of the 1952–54 british north greenland expedition from their temporary base at kap rink, and in 1980 by a group from exercise icy mountains vi. (mt. matterhorn.) matterhorn ‘s’ peak 73ø (73°25.8´n 27°36.3´w). mountain on the south side of jættegletscher, the present lille cervin, so named informally by louise boyd’s 1931 expedition. it appears on some of boyd’s maps marked ‘s’. mattmarksø 73ø-311 (73°51.3´n 23°16.8´w). lake in central hudson land. named during lauge koch’s 1936–38 expeditions by heinrich bütler after the mattmarksee in vispertal, switzer land. maud sø 73ø-680 (73°35.8´n 26°57.1´w; map 4). lake in west andrée land. named by john haller following explorations during lauge koch’s 1949–51 expeditions, after queen maud [1869– 1939], a daughter of edward vii of great britain who married haakon vii of norway. (maud-see.) mauritius tinde 73ø-375 (73°41.7´n 26°24.7´w). mountain in north andrée land, south of eremitdal, named by erdhardt fränkl during lauge koch’s 1948–50 expeditions. (mauritiustinde.) mauritz diesens sjö 73ø (73°43.7´n 24°40.4´w). lowest and largest lake in brogetdalen in strindberg land, the present laksesø. the name is only used by munsterhjelm (1937), and was named after mauritz diesen, a norwegian lawyer who fished here with munsterhjelm in 1936. maursundet 73ø (73°03.3´n 23°04.9´w). sound between robertson ø and north geographical society ø. used only on nsiu maps (lacmann 1937), the name is a norwegian expression for a sound with a strong current. mausa 74ø (74°09.5´n 20°36.3´w). river on se clavering ø, the present moskusokseelv. so named on the nsiu maps of lacmann (1937) after the mauseeidvåg in the møre and romsdal district of norway, home of the norwegian hunter peder røbek (see also røbekfjellet). mausdalen 74ø (74°09.5´n 20°36.3´w). valley on se clavering ø containing the river mausa, equivalent to the present baesdalen. the name is used on an nsiu map (1932a). mckenzie glacier – note that ‘mc’ is treated as ‘mac’. mears fjeld 71ø (71°56.6´n 25°12.4´w; map 5). peak 2100 m high in the upper reaches of sefström gletscher, stauning alper. climbed by the 2001 scottish mountaineering club expedition. medalselva 73ø (73°34.0´n 21°41.0´w). river flowing into macken zie bugt, so named on an nsiu map (1932a; fig. 13). derived possibly from a similar name in the jotunheim area of norway. medusagryde 72ø-387 (72°02.1´n 23°21.5´w). bowl-shaped valley with a small glacier on the east side of majdal, north scoresby land. named by hans kapp during lauge koch’s 1957–58 expeditions, for the shape, somewhat reminiscent of a jelly fish (= medusa). mehrendalen 74ø (74°02.1´n 22°52.8´w). valley in north hudson land draining north to wordie gletscher, equivalent to the present slugtdalen. so named on the nsiu maps of lacmann (1937) after martin mehren [b. 1905], a norwegian who, with arne høygaard, made a crossing of greenland from west to east in 1931. 252 melander river 72ø (72°31.9´n 23°54.5´w). name used by 1968–74 dundee university expeditions for the river in eskdal, sw traill ø, which drains into karupelv. melch dal 72ø-454 (72°53.7´n 26°49.1´w; map 4). valley on the north side of dickson fjord from which a conspicuous, white, foaming waterfall drains south into the fjord. it was named during the 1931–34 treårsekspeditionen by eugène wegmann, after melchtal north of lausanne, switzerland. mellemdal 74ø-378 (74°42.2´n 22°13.6´w). valley joining tyroler dal and svejstrup dal. the name was adapted from the verbin dungs tal (= connecting valley) of mittelholzer (1941), at the sug gestion of w.r.b. battle in 1948 (mellemdal = between valley). mellemfjeld 75ø (75°10.2´n 19°50.6´w). name occasionally used by danish hunters in the 1930s for a hill between their kystfjæld (søndre muschelbjerg) and nordre muschelbjerg (nyholmpoulsen 1985). mellemfortet 78ø-26 (78°23.5´n 19°41.2´w; map 4). island in jøkelbugten, east of nørre mellemland. named during the 1938– 39 mørkefjord expedition after the island fortress of the same name off copenhagen, where the expedition’s dogs were housed in transit. the name was said to continue the tradition of naming features in the region after copenhagen locations, that was begun by the 1906–08 danmark-ekspeditionen. mellemgletscher 71ø-243 72ø-300a (72°00.3´n 24°04.9´w; map 5). the middle of three glaciers draining into the head of deltadal, north werner bjerge. the name originated from a climbing excursion during lauge koch’s 1950 expedition (styger 1951). (mellemgletscher.) mellemhuset – see also midtstua. mellemhuset 71ø (c. 71°46´n 22°57´w). norwegian hunting hut built in 1932–33 for helge ingstad’s expedition at the mouth of solfaldsdal, about halfway along fleming fjord. no trace of it remains (p.s. mikkelsen 2008). it was also known as syveren, pasdalshuset and funkis. (mellem-huset, midthuset.) mellempas 71ø-242 (71°59.4´n 24°10.7´w; map 5). pass between the heads of mellem gletscher and arcturus gletscher, werner bjerge. named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk. melles lake 76ø (76°07.7´n 18°37.9´w). lake on store koldewey where sampling was undertaken for phytoplankton studies (cre mer et al. 2005). menageløsdal 70ø (70°15.0´n 27°30.0´w). e–w-trending valley in central gåseland, draining east to gåsefjord. the name is thought to have arisen with the prospecting teams of nordisk mineselskab in the late 1960s, and to be a reference to the apparent absence of any animal life (musk ox, hares). the name was used as a reference locality by larsen et al. (1989). menander spir 72ø-495 (72°19.1´n 24°31.2´w; map 5). sharp rock summit 1622 m high in the syltoppene overlooking the menander øer. climbed by the cambridge university expedition on 11 august 1963. menander øer [immikkeerterajii] 72ø-23 (72°20.6´n 24°17.4´w; maps 4, 5). line of several small islands on the sw side of kong oscar fjord. named menanders öar by a.g. nathorst’s 1899 expedition after j. menander, 2nd mate of the antarctic, the expedition ship. (menander islands, menanderöyane.) menanders bugt 72ø (72°30.5´n 24°04.7´w). name occasionally used for the present holm bugt, north of the menander øer (e.g. hansen 1982). méneset 72ø (72°42.3´n 22°42.9´w; fig. 14). peninsula on south geographical society ø, east of silja ø. used only on nsiu maps (lacmann 1937), and so named because the peninsula is a convenient point on which to make a bearing (mé = bearing). (meneset.) mercanton gletscher 73ø-552 (73°00.0´n 27°54.0´w). glacier in goodenough land, draining south to join charpentier gletscher, named by j.m. wordie’s 1929 expedition as mercanton glacier. usaf aeronautical charts show it draining north to nordenskiöld gletscher. see also mercantonbreen. (mercantons gletscher.) mercantonbreen 74ø (74°11.0´n 22°25.8´w). lobe of wordie gletscher between scotstounhill and jordanhill. so named on the nsiu maps of lacmann (1937) after paul louis mercanton [1876–1963], a swiss meteorologist and glaciologist noted for his work on swiss glaciers, and on international commissions. he accompanied de quervain on his crossing of the greenland ice cap in 1912. merchiston tinde 72ø-371 (72°04.9´n 24°48.3´w; map 5). mas sive mountain peak 2400 m high near the head of bersærkerbræ, north stauning alper. first climbed by malcolm slesser’s 1958 expedition, and named after the castle near edinburgh. sub se quent ascents were made in 1963, 1968 and 1969. (merchiston tinde.) mercurius gletscher 71ø-339 (71°39.1´n 25°03.0´w; map 5). glacier in the south stauning alper, flowing east to join bjørnbo gletscher. named mercury glacier by john hunt’s 1960 expedition, after the planet mercury. mercury passet 71ø (71°37.5´n 25°13.7´w; map 5). pass between the fig. 59. view northwards from femdalen across ardencaple fjord to dronning margrethe ii land, with matterhorn and the barth bjerge to the right. the john haller photograph collection, geus archive. dronning margrethe ii land matterhorn barth bjerge ardencaple fjord femdalen 253 upper parts of mercurius gletscher and oxford gletscher, south stauning alper, first traversed by the 1970 dundee university expedition. (mercurius passet.) meridiannunatak 76ø (76°25.7´n 22°37.0´w). dark peak east of l. bistrup bræ. the name was introduced by j.p. koch and alfred wegener during their 1912–13 expedition, and refers to a pointed peak used as a surveying mark about 30 km south of their wintering station, probably in westernmost rechnitzer land. merthyr peak 70ø (70°50.8´n 26°04.6´w). summit on the north side of korridoren, milne land, reached from the south by a narrow ridge of crumbly rock. climbed by the 2004 west lancashire scouts expedition. messterfjellet 74ø (74°16.5´n 21°03.9´w). mountain 1308 m high on central clavering ø. the name is used only on nsiu maps (lac mann 1937), and was given for eduard o. messter [b. 1893], a director of zeiss aerotopograph gesellschaft jena. mesters vig 72ø-20 (72°08.3´n 23°47.3´w; maps 4, 5). deep bay or small fjord on the sw side of kong oscar fjord. named by a.g. nathorst’s 1899 expedition as mästers vik. the assumption that it was given for the master of the antarctic, the expedition ship (see forsblad fjord), is probably incorrect. as the swedish term ‘mäster’ is not synonymous. svend sølver (personal communication 2003) suggests it was more likely intended to commemorate the chief engineer (maskinmester), i. peterson. the name mesters vig has also been commonly used for the airfield west of the bay. see also mestersvig. (masters bay, mästerbukta, mestersvig fjord.) mesters vig flyveplads – see mestersvig. mesters vig glacier 72ø (72°05.5´n 23°55.5´w). term used by pessl (1962) for the glacier formerly occupying deltadal and mesters vig. mestersvig 72ø-20a (72°13.9´n 23°55.1´w; maps 3–5). airfield north of mesters vig, west of noret, constructed in 1952 as part of the government agreement with nordisk mineselskab concerning the exploitation of lead at blyklippen (p.s. mikkelsen 2005). the gravel runway is 1800 m long. additional buildings were added during the 1950s and 1960s, the last major addition being a radio station and control tower erected in 1977–79. the airfield was officially closed on 15 october 1985, but continues in use, maintenance being carried out by a small military group, ‘forsvarets vagt mestersvig’. many of the main buildings have been given names (millionæren, hilton, olympos, blåtårn, rødull, valhal, havne bygningen, luftkastellet; see p.s. mikkelsen 1994, 2008), but these are not listed in this catalogue. the name mestersvig (in one word) was not officially approved until the late 1980s, but has been very commonly used as a designation for the airfield in official and unofficial documents since its construction. in the earliest days of its existence it was sometimes referred to as government station or flyverplads (washburn 1965). until the airfield at constable pynt came into service it was the principal airfield in this part of east greenland used by visiting expeditions and also served the settlement at scoresbysund. (mestersvig station, mesters vig flyveplads.) metacarpal 72ø-514 (72°01.5´n 25°21.9´w). mountain on the sw side of sefström gletscher, very close to inverarnan. named by the 1963 cambridge university expedition. metaforgletscher 76ø-338 (76°15.9´n 26°09.5´w; map 4). small glacier in sw dronning louise land, flowing sw to join ebbe gletscher. the name was given by the british north greenland expedition 1952–54, and was derived from the greenlandic custom of using picture language (metaphor) in giving names. meydenbauerfjellet 72ø (72°55.1´n 22°51.6´w). mountain on central geographical society ø, the present knolden. so named on the nsiu maps of lacmann (1937) after a. meydenbauer [1834– 1922], a german architect and archaeologist. meyer-hus 74ø (c. 74°28´n 21°03´w). norwegian hunting hut on the north side of tyrolerfjord, south of zackenberg. built by the foldvik expedition in september 1927, and named after meyer olsen, a norwegian hunter who helped build it. it has also been known as trangfjordhuset and zackenberghuset. exact location un known (p.s. mikkelsen 2008). meyerstein bjerg 75ø-25 (75°18.0´n 17°57.0´w). mountain 305 m high on ne shannon. named by karl koldewey’s 1869–70 expedition as meyerstein berg, probably after moritz meyerstein [1808– 1882], an instrument maker in göttingen, who supplied meteorological and surveying instruments to the expedition (j. løve, personal communication 2010). (meyersteins bjerg.) miami fjeld 76ø (76°10.3´n 18°40.0). mountain north of træk passet, store koldewey. the name is used on 1952 ams maps, and is also found in den grønlandske lods (1968) in the form miami bjerg. micardbu 77ø-111 (77°04.3´n 18°11.4´w; map 4). norwegian scientific and hunting station 5 km north of fyrretyve kilo meter næsset on the east coast of germania land, of which only the foundations now remain. the remains of the house were taken down in 1960 and used to build a smaller hut for danmarkshavn weather station. named after the leader of the norsk–fransk polar ekspedisjon 1938–39 which had built the station. count gaston micard [1879–1961], an eccentric frenchman, spent several summers and winters in east greenland waters using chartered norwegian sealers, and was noted for always sheltering under a yellow silk umbrella patterned with streaks and blotches (knutsen 1949). he was one of the original, large share owners of the suez canal. micard was taken ill during the winter of 1938–39, and evacuated by a stinson seaplane operating from the veslekari. michelangelo kløft 77ø-92 (c. 77°10´n 19°32´w). river gorge leading down from slædelandet to fladebugt in skærfjorden. so named during the 1938–39 mørkefjord expedition because of a 10–12 m high rock resembling the statue of a man, whose shape, posture and pathos was to eigil knuth reminiscent of a roughly made statue by michelangelo. (michelangelos kløft, michelangelo kløften.) middelgrunden 75ø (75°58.4´n 20°10.3´w). island in the mouth of grandjean fjord, the present trums ø. this name appears only on the sketch map by t. johansen published in koch (1940), a map drawn during the original exploration of the fjord in 1932. it is a common danish name for an offshore shoal area, and was perhaps given for the small island of the same name in the mouth of fur sund, jylland, or the fort off copenhagen. middle gneisnæs 76ø (76°14.3´n 18°34.3´w). name used as a geological reference locality by frebold (1935) and maync (1949) for a point on the east coast of store koldewey between nordre gneisnæs and sydlige gneisnæs. (mittlere gneisnaes.) middle peak 72ø (c. 72°08´n 25°03´w). peak in the north stauning alper, climbed by g. dionisi’s 1982 expedition during a traverse including norsketinden. midnatspas 73ø-427 (73°21.8´n 24°43.8´w). pass across the west part of gunnar andersson land, ymer ø. this name was given by a.b. cleaves and e.f. fox during john k. howard’s 1933 expedition, because their geological work began at midnight after they had made a traverse of the pass. the name was adopted by the next geologist to work in the region (eha 1953), and approved in its present form. midnight peak 71ø (71°38.5´n 25°09.5´w; map 5). peak about 1700 m high on the south side of mercurius gletscher, south stauning alper. first climbed by james clarkson’s 1961 expedition, and so named because they reached the summit at exactly midnight. midnight sun 70ø (70°47.0´n 22°03.9´w). summit 930 m high in liverpool land, north of bjerring pedersen gletscher. the name was recorded by the 2002 loughborough grammar school expedition that made the second ascent. midterfjeld 73ø-368 (73°39.4´n 24°43.3´w; map 4). mountain about 1200 m high in south strindberg land, named by hans r. katz during lauge koch’s 1948–49 expeditions (midter = middle). 254 midterfjellet 73ø-155 (73°29.8´n 20°32.0´w). mountain 752 m high in se hold with hope, named for its position between two other peaks. an nsiu map (1932a) used the name méfjellet, while gustav thostrup used kommafjeldet for the same feature. midterholmen 76ø-290 (76°43.8´n 20°46.9´w; map 4). island in the west part of dove bugt. so named by the 1938–39 mørkefjord expedition because of its position centrally between ringøen and rødeø. anthons ø has also been used. midternæs 77ø-88 (77°05.0´n 20°46.1´w). cape almost in the middle of the north shore of sælsøen. named by the 1938–39 mørkefjord expedition. midternæshytten 77ø (c. 77°05´n 20°48´w). danish hunting hut at midternæs on the north side of sælsøen. built by nanok in novem ber 1938, it has now disappeared. according to p.s. mikkelsen (1994) it is identical with inderhytten. midtstua 73ø (73°48.4´n 21°45.8´w). norwegian hunting hut on the east coast of loch fyne, south of strømmen and about 10 km north of herja elv, built by the foldvik expedition in august 1926. it was named for its position halfway along the fjord. rebuilt in 1954, it is now more or less a ruin. (mellemhuset, midthuset.) midthuset – see midtstua (hut east of loch fyne) and syveren (hut in fleming inlet). midway nunatak 69ø (69°07.6´n 32°44.4´w). reference name used for a nunatak in the prinsen af wales bjerge, northern kong christian ix land (nielsen et al. 2001). mikael bjerg 71ø-58 (71°09.5´n 23°05.1´w). mountain in eastern jameson land. named during lauge koch’s 1926–27 expeditions by alfred rosenkrantz and tom harris as mt mikael or mt mikael fjæld after their greenlandic assistant mikael kunak. mikkel sø 70ø-373a (70°52.9´n 22°28.6´w). small lake at the south end of klitdal. named by svend funder, who carried out borings in the lake during ggu expeditions in the 1970s. it was named for the fox dens by the lake; in denmark ‘mikkel’ is a common nickname for a fox. milano gletscher 70ø-274 (70°03.9´n 23°00.0´w; map 4). glacier on volquaart boon kyst. it was first explored by leonardo bonzi’s 1934 expedition, which named it ghiacciaio milano after the town of milan, the expedition’s starting point. milepælen 78ø-28 (78°37.1´n 23°08.5´w; map 4). southernmost and highest peak of moltke nunatak. the name was suggested by the place name committee as a substitute for a proposal by the 1938–39 mørkefjord expedition. the peak was climbed by svend sølver on 27 may 1939. (milepæl = milepost). milne land 70ø-23 (70°43´n 26°48´w; maps 3, 4; figs 7, 56). large island bounded by hall bredning, fønfjord, røde fjord and øfjord. named by william scoresby jr. in 1822 as a compliment to sir david milne [1763–1845], who had a long and distinguished career in the royal navy. (milnes land, terre de milne, milneland.) mimingbreen 74ø (74°13.5´n 20°48.4´w). glacier on se clavering ø. so named on the nsiu maps of lacmann (1937) after miming, a wood troll in old nordic mythology. minebyen 72ø (72°11.8´n 24°05.9´w). name commonly used for the mining town in store blydal which was operated by nordisk mine selskab between 1956 and 1962. the town consisted of a large number of barracks and workshops, a few of which were moved to nyhavn in the 1970s. at the height of activities 150 persons were employed in the summers and about 50 in the winters. in the 1980s the site was gradually cleared, and by 1988 only two houses remained. mineralbugten 72ø (72°53.7´n 25°42.1´w). bay on the coast of se suess land, the present lumskebugten. the name arose during the 1931–34 treårsekspeditionen and is found in a number of contemporary reports, e.g. th. sørensen (1933). origin of name uncertain, but it may refer to local mineralisations. (mineral bay.) mineralbukta 72ø (72°53.8´n 25°43.9´w). norwegian hunting hut on the west side of the floodplain at the head of lumskebugten (also known as mineralbugten), built by arktisk næringsdrift in september 1934. it has also been known as lumskebugthytten and sunnmøresheimen. mineralsee 71ø (71°46.9´n 23°56.6´w). lake on the south side of mineralseepingo, in pingo dal, north jameson land. the name was used by fritz müller during lauge koch’s 1954–55 expeditions, and was given for the calcareous and gypsum-rich waters. mineralseepingo 71ø (71°46.9´n 23°56.6´w). name used by müller (1959) in his report on work during lauge koch’s 1954–55 expeditions, for an 18 m high pingo in pingo dal. see also mineralsee. minger bjerg 72ø-455 (72°58.2´n 27°04.8´w; map 4). ice plateau in sw suess land, climbed by eugéne wegmann in august 1933 during the 1931–34 treårsekspeditionen. it was named after the mingerhubel, an ice plateau in switzerland. minimalen 71ø (71°53.1´n 23°00.9´w). norwegian hunting hut in henrik møller dal, built by helge ingstad’s 1932–34 expedition. the name refers to the small size of the hut, which was originally known as øyedalshytten. it is now a ruin. minus four 70ø (70°54.0´n 25°56.0´w). summit about 1800 m high on northern milne land. it was climbed by the 1989 greenland milne land expedition. mirakeldal 72ø-378 (72°01.1´n 23°39.1´w). small valley on the south flank of oksehorn, draining into kolledalen, north scoresby land. so named by hans kapp during lauge koch’s 1957–58 expeditions, because the valley allegedly gave rise to miraculous geological discoveries. misanthropen fjelde 79ø-43 (79°11.4´n 20°03.6´w). snowcapped mountain in lambert land. the name is one of a group of five given by the place name committee for dogs used on the 1906–08 danmark-ekspeditionen. they replaced names suggested by john haller. ‘misanthropen’ was an old and rather miserable dog which did not get on with the other dogs in the team. missing ring 71ø (71°13.1´n 26°33.0´w). high point (2110 m) on the ice cap north of edward bailey gletscher, renland. climbed and named by the 2007 west lancashire mountaineering group expedition. misteltengletscher 74ø-380 (74°21.6´n 20°54.1´w). glacier on north clavering ø. originally named on 1937 nsiu maps in the form mistelteinbreen, after an enchanted sword of old nordic mythology, mistelten (made of mistletoe), with which høder killed balder. the name was first approved for general usage in 1950. mittenwalder tinde 71ø (71°50.1´n 25°30.2´w; map 5). peak about 2500 m high on the south side of the upper basin of spærre gletscher, climbed by karl m. herligkoffer’s 1966 expedition on 18 august. two of the climbers, michl anderl and gebhard plangger, were mountain guides in mittenwald/luttensee. mitternachtspitze 72ø (72°03.9´n 25°40.9´w). mountain se of trekant at the head of trekantgletscher. it was climbed and so named by wolfgang weinzierl’s 1970 expedition. this peak is probably identical with damslottet. (midnight peak.) mitterspids 72ø (72°01.2´n 25°04.2´w; map 5). mountain on the north side of sefström gletscher, stauning alper, where kirkbrae and a smaller unnamed glacier meet. named and first climbed by hans gsellman’s 1957 expedition. (mitterspitze.) mittlandet 73ø (73°33.0´n 27°54.0´w). name used by anrick (1932) for the landmass between gerard de geer gletscher and jætte gletscher, the present louise boyd land. mittleres profil 74ø (74°43.8´n 20°01.9´w). geological reference locality on se kuhn ø, used by maync (1947) in his description of work during lauge koch’s 1936–38 expeditions. mobu dal 73ø-305 (73°38.1´n 23°27.3´w; map 4). valley on west gauss halvø, running ne to moskusoksefjord. the name was used by swedish geologists during the 1931–34 treårsekspeditionen, and is said to be an abbreviated name of a geological society. modiolaelv 70ø-145 (70°32.5´n 22°40.6´w). minor river in se 255 jameson land running into lakseelv nw of kap stewart. named by alfred rosenkrantz during lauge koch’s 1926–27 expeditions as modiola elv after the fossil lamellibranchs. moena tinde 72ø (c. 72°06´n 25°15´w). peak 1940 m high on the south side of gully gletscher, north stauning alper. climbed on 2 august by sandro pucci’s 1984 expedition. origin of name unknown. molehill 71ø (71°55.0´n 24°58.6´w). small peak about 2300 m high at the head of dalmore glacier, central stauning alper. so named by the 1968 dundee university expedition, which made the first ascent. (the molehill.) molen 70ø-126 (70°52.8´n 22°43.9´w). mountain nw of the head of hurry inlet, named by tom harris and alfred rosenkrantz during lauge koch’s 1926–27 expeditions in the form mole mountain. mollytinde 71ø (71°59.4´n 24°50.8´w; map 5). low mountain close to camp on the west side of storgletscher, only 1670 m high, central stauning alper. climbed and named after a living person by the 2007 smc east greenland expedition. mols bjerge 72ø-81b (72°29.6´n 22°28.3´w). mountain range north of mountnorris fjord, east traill ø. named by ove simon sen during the 1931–34 treårsekspeditionen after the danish locality of the same name in jylland. moltke bjerg 74ø-71 (74°24.8´n 20°50.8´w). mountain 1388 m high on north clavering ø. named by karl koldewey’s 1869–70 expedition as berg moltke, after count helmuth karl bernhard von moltke [1800–91], field marshal in the german army, and one of those present when the expedition sailed from bremerhaven in june 1869. (moltkefjellet, moltkes bjerg.) moltke nunatakker 78ø-20 (78°37.3´n 23°00.0´w; map 1). nunatak group west of hertugen af orléans land. named by the 1909–12 alabama expedition as moltke’s nunatakker after count carl poul oscar moltke [1869–1935]. moltke was a danish naval officer and had participated in vilhelm garde’s 1893 expedition to sw greenland, and led an expedition to the julianehåb region in 1894. he later followed a career in the diplomatic service and was foreign minister from 1924 to 1926 (j. løve, personal communication 2009). moltvika 72ø (72°37.9´n 22°43.3´w). bay on the ne side of traill ø, west of kap palander. used on the nsiu maps of lacmann (1937), the name was given for walter molt [b. 1901], a norwegian hunter who wintered in east greenland in 1932–33 and 1934–35. mona bjerg 73ø-713 (73°09.6´n 28°17.9´w; see also fig. 65). moun tain c. 2300 m high ne of nordenskiöld gletscher, named by noel ewart odell after his wife (odell 1937a). gwladys mona odell [1891–1977] accompanied her husband on many of his expeditions between 1919 and 1961. in 1933 they had climbed mona bjerg together during louise boyd’s expedition. the name (originally mount mona) was revived, and approved, in 1952 at the suggestion of john haller. monacleus 74ø-128 (74°12.0´n 20°49.1´w). mountain on se clav ering ø with one summit 1403 m high. named during lauge koch’s 1929–30 expeditions by h.g. backlund in the form mt. monacleus. see also binucleus and trinucleus. (monaclius bjerg.) mont bertram 70ø (70°41.9´n 25°58.8´w). mountain 1300 m high on se milne land on the south side of charcot gletscher. so named in the report by parat & drach (1934) on their work with j.-b. charcot’s 1933 expedition, after one of the members of the 1933 cambridge expedition which was transported to and from greenland by the pourquoi pas?. g.l.c. bertram had worked on bjørneøya in 1932, in graham land (antarctica) from 1934 to 1937, and subsequently in the middle east. he was director of the scott polar research institute from 1949 to 1956. mont blanc de furesoe 71ø (71°53.7´n 25°54.8´w). highest point on the ice cap on the west side of prinsessegletscher, south of furesø (2570 m). named and first climbed by claude rey’s 1968 expedition. mont du pollux 70ø (70°42.1´n 26°02.9´w). mountain 1510 m high on se milne land, nw of bay fjelde. the name was used during j.-b. charcot’s 1933 expedition (parat & drach 1934), and was named after the french icebreaker pollux that accompanied the pourquoi pas? on the 1932 and 1933 expeditions. mont freudo 71ø (71°52.8´n 25°43.3´w; map 5). mountain about 2480 m high on the east side of prinsessegletscher, south of furesø. named and first climbed by claude rey’s 1968 expedition. mont lack 70ø (70°45.5´n 26°01.9´w). mountain 1720 m high on the north side of charcot gletscher, east milne land. so named in the report on j.-b. charcot’s 1933 expedition by parat & drach (1934). named after david lambert lack [1910–73], a member of the 1933 cambridge expedition transported to and from greenland with the pourquoi pas? lack was a noted ornithologist, who from 1945 was director of the edward grey institute of field ornithology. mont rosenkrantz 70ø (70°40.3´n 25°51.0´w). mountain 1028 m high on east milne land equivalent to the present pourquoi pas tinde. the name was used during j.-b. charcot’s 1933 expedition (parat & drach 1934), and was given for alfred rosenkrantz [1898–1974], a noted danish geologist who had been the first to investigate the rocks of this part of milne land (rosenkrantz in: koch 1929). rosenkrantz worked in east greenland, initially with lauge koch’s expeditions, between 1926 and 1936, and subsequently led numerous expeditions to the nuussuaq–svartenhuk area of west greenland. he was professor of geology at the uni versity of copenhagen from 1953 to 1966, and one of the prime instigators of the geological survey of greenland. mont röhling 74ø (74°12.4´n 20°55.9´w). mountain c. 1434 m high on south clavering ø. named by lauge koch’s 1929–30 expeditions after frederik holger røhling, a superintendent in the tech nical department of the geodetic institute, who was an expert in drawing and lithographic techniques. the name was used in several geological publications, and appears on lacmann’s (1937) maps, but was never approved. (röhlingfjellet.) mont saussure 71ø (71°55.2´n 25°23.5´w). mountain 2580 m high on the north side of duart gletscher, stauning alper, now known as duart borg. it was first climbed by the 1964 aac zürich expedition, which named it after h.b. de saussure, a geneva scientist. see also saussure massiv. the second ascent was made by karl m. herligkoffer’s 1966 expedition, which called it dresdner spids. mont wegener 70ø (c. 70°40´n 21°59´w). mountain about 1400 m high in south liverpool land, probably the highest peak of tvil lingerne, or possibly korsbjerg. the name is used in a report by rothé (1941) on the work at the french international polar year station at scoresbysund, and was given for alfred lothar wegener. see also wegener halvø. montane 71ø (71°15.6´n 26°14.4´w). point 2201 m high on the ice cap north of edward bailey gletscher, renland. climbed and named by the 2007 west lancashire mountaineering group expedition. monte bello 73ø-389 (73°37.2´n 26°00.0´w). mountain in andrée land. named during lauge koch’s 1948–50 expeditions by erd hardt fränkl, it translates as beautiful mountain, and is a com mon place name (e.g. the castle in bellinzona, southern switzer land is known as montebello). monte carmela 72ø (72°20.5´n 24°43.2´w). mountain in the north stauning alper, sw of kap peterséns. location uncertain, but probably the 1589 m high peak blåhorn. named during the 1930 nsiu expedition by vittorio beonio-brocchieri, who climbed the mountain in august 1930, and named it after his mother. see also broc chieridalen. monte somma 71ø-282 (71°54.1´n 23°52.3´w; map 5). mountain 1885 m high in the south werner bjerge. named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk, 256 and climbed by wenk in 1953. it apparently resembles monte som ma on vesuvias. monts aldinger 70ø (70°41.3´n 26°06.1´w). range of mountains on se milne land, north of vinkeldal, up to 1620 m high. so named in the report by parat & drach (1934) describing work during j.-b. charcot’s 1933 expedition, to commemorate herman aldinger, a geologist who worked in this region in 1933. see also aldinger elv. monumentet 76ø-331 (76°27.7´n 25°04.2´n; map 4). prominent mountain south of pony gletscher in dronning louise land. named by the british north greenland expedition 1952–54 (monu mentet = the monument). monumentet – see danmarks monumentet. moorsom dal 72ø-352 (72°10.3´n 22°09.7´w). valley on se traill ø. named by h.p. heres during lauge koch’s 1956–58 expeditions for its proximity to kap moorsom. moro bjerg 73ø-688 (73°26.1´n 26°04.2´w). mountain in south andrée land. named by john haller following explorations during lauge koch’s 1949–51 expeditions, for a resemblance to monte moro on the swiss–italian border south of saas fee. morris bjerg 72ø-238 (72°18.0´n 22°57.5´w). mountain 942 m high on se traill ø, west of steenstrup bjerg. named by desmond t. donovan during lauge koch’s 1949–50 expedition after john morris, a 19th century palaeontologist who worked on fossils of the same age as the rocks which make up the mountain. morten sø 70ø-372a (70°53.0´n 22°26.9´w). small lake at the south end of klitdal. named by svend funder who made borings in the lake during a ggu expedition in the 1970s. ‘morten’ is a danish name sometimes used for ‘martin’, a nickname commonly used for the goose, the traditional dish eaten on ‘mortensaften’ on 10 november. morænebakkerne 74ø (74°30.8´n 20°36.6´w). moraine deposits at the east end of store sødal, near zackenberg. the name is used as a reference locality by scientists at zackenberg forskningsstation (meltofte & thing 1996). morænedal 73ø-650 (73°41.2´n 25°09.8´w). valley in ne andrée land, draining into geologfjord. so named during the 1931–34 treårsekspeditionen by th. johansen, because it is a typical glacial valley with moraine ramparts. morænedalshytten 73ø (73°41.3´n 25°06.2´w). norwegian hunting hut built in 1938 for arktisk næringsdrift on the north side of morænedal. it is also known as hastværkshytten. now a ruin. morænelandskab 77ø (c. 77°30´n 21°34´w; map 4). area of spectacular moraines on the east side of kofoed-hansen bræ. used in this form by the 1906–08 danmark-ekspeditionen, but probably in tend ed as a descriptive rather than a formal name. a photograph appears in koch (1912). (moränenlandschaft.) morænepynt 70ø-19 (70°27.5´n 27°26.1´w). small peninsula on the south coast of milne land. so named by carl ryder’s 1891–92 expedition because of the gravel deposits (moraine) and fossilbearing clays found here. the name hvalpynten is used in helge vedel’s diaries of carl ryder’s 1891–92 expedition; hval = whale (gulløv 1991; j. løve, personal communication 2010). (moræne pt.) morænevolden 77ø-42 (77°18.1´n 23°39.8´w). moraine ridge north of søstersøer, encountered by vilhelm laub in northern dronning louise land during the 1909–12 alabama expedition. it was probably intended as a descriptive rather than a formal name. moræneø 76ø-286 (76°25.9´n 21°46.6´w). small island at the entrance to bræfjorden, at the foot of a glacier descending from rechnitzer land. so named by the 1938–39 mørkefjord expedition because it consists of gravel deposited by the glacier (= moraine). mosaikskærene 76ø-287 (76°48.1´n 20°47.4´w). skerries near spydodden, off eastern daniel bruun land. so named by the 1938– 39 mørkefjord expedition because of the red and white mosaic-like patterns in the rocks. moschusochsenalm 75ø (c. 75°19´n 17°50´w). plain rising northwards from the base camp of the 1943–44 operation bassgeiger at kap sussi, where six musk ox were seen in mid-february 1944, of which one was shot. the name is reported by olsen (1965). mosen 76ø-254 (76°48.7´n 19°03.3´´w). name used in the ornithology reports of the 1906–08 danmark-ekspeditionen for a part of winge kyst near stormkap, south germania land, where there are numerous small lakes. moskusdalen 71ø (71°55.8´n 23°58.6´w). name given by the 1930– 32 møre expedition to the first side valley to blomsterdal, south of their hunting station at antarctic havn. it is possibly the present flexurdal (rogne 1981). named for the numerous musk oxen. steinrøisdal has been used for the same valley. moskusdalen 73ø (73°55.8´n 23°58.6´w). name used by sigurd skaun and harald welde in 1932 for the present johan davidsen dal, the boundary between ole rømer land and west hudson land. it was named for the numerous musk ox (fig. 60). moskusfjorden 76ø (76°49.6´n 19°27.5´w). name used during the 1906–08 danmark-ekspeditionen for the locality where a.l.v. manniche had shot musk oxen in 1906 (poulsen 1991), probably the bay east of snenæs. moskusheimen 74ø (74°21.8´n 21°51.7´w). norwegian hunting station on the west side of the sound revet, west of clavering ø. built by henry rudi during the devold expedition in 1928, and so named because while building it a flock of musk ox came down from the hills and milled around the hut as if trying to enter it. a smaller hut nearby built by the foldvik expedition in 1927 subsequently served as a depot house. the nsiu list of huts by orvin (1930) indicates this hut as tyrolerheim, a name also applied to two other huts. this station has commonly been referred to after its location as revet, which is the approved name. (moskus-heimen.) moskusokseelv 70ø-137 (70°38.8´n 22°40.9´w). river in moskus okse kløft on the west side of hurry inlet. the name was first used in the report by harris (1931) on his work during lauge koch’s 1926–27 expeditions as musk ox river. the same name was used in error by roberts (1935) for the present gåseelv. moskusokseelv 74ø-101 (74°09.5´n 20°36.3´w). river on se clavering ø, named during lauge koch’s 1929–30 expeditions in the form muskox river or musk-ox river, after the numerous musk oxen. it has also been called mausa and giskovselv. (muskusoksen elv.) moskusoksefjeld 73ø-370 (73°37.3´n 24°41.5´w; map 4). moun tain about 1100 m high in south strindberg land, named by hans r. katz during lauge koch’s 1948–49 expeditions after the musk ox. (moskusoksefjæld.) moskusoksefjeldene 76ø-32 (76°55.3´n 19°30.2´w). range of hills in germania land east of hvalrosodden, named by the 1906–08 danmark-ekspeditionen. traces of musk ox were seen nearly everywhere by the expedition, and some were shot here. (moskus fjældene, musk-ox-mountains, moskusoksefjelde, muskox mts.) moskusoksefjord 73ø-32 (73°40.0´n 22°20.0´w; maps 2, 3, 4). fjord between moskusokselandet and gauss halvø, named by a.g. nathorst’s 1899 expedition as myskoxefjorden because he saw 67 musk oxen on the fjord sides on his first exploration of the fjord. (muskoxen fjord, muskox fiord, moskusoxe fjord, moschusochsen fjord, musk-ox fiord, moschusochs fjord, moskusfjorden.) moskusoksehytte 73ø (73°33.3´n 20°30.5´w). name for a hut in se hold with hope built in august 1927 by the foldvik expedition, which was partly covered by musk-ox skins. the hut has also been known under the names bukta, tvivlsom and skandalen. (moskus hytte.) moskusoksekløft 70ø-136 (70°38.8´n 22°40.9´w). ravine on the west side of hurry inlet, named by alfred rosenkrantz during lauge koch’s 1926–27 expeditions in the form moskusokse kløft, after the numerous musk oxen. (musk-ox kløft.) moskusokselandet 73ø-33 (73°45.0´n 23°15.0´w; map 4). sw part of hudson land, between moskusoksefjord to the south, and 257 ankerbjergselv, visp and johan davidsen dal to the north. named by a.g. nathorst’s 1899 expedition as myskoxelandet after the abundant musk oxen seen on the slopes leading down to moskus oksefjord. the present usage is more restricted than nathorst’s, and corresponds more or less to that used by seidenfaden (1931), who employed the term musk-ox range or moskusokse bjærge. (muskoxen land, moschusochs gebirge, muskusoksefjella.) moskusoksesø 73ø-369 (73°37.9´n 24°50.2´w). lake in south strindberg land, west of moskusoksefjeld. named during lauge koch’s 1948–49 expeditions by hans r. katz after the musk ox (fig. 60). moskusøyra 74ø (74°21.3´n 21°50.8´w). large delta on west clav ering ø, the present tangen. used only on nsiu maps (lacmann 1937), and named after the numerous musk oxen seen here. mosquito ridge 70ø (70°33.8´n 22°54.7´w). ridge on the west side of møns elv, southern jameson land. named by herman aldinger during the 1931–34 treårsekspeditionen after the abundant mos qui toes. mount brassica 71ø (71°13.1´n 26°21.7´w). point 2065 m high on the ice cap north of edward bailey gletscher, renland. climbed and named by the 2007 west lancashire mountaineering group expedition. mount fatigue 70ø (70°47.8´n 26°06.4´w). summit on the south side of korridoren, milne land. climbed by the 2004 west lan cashire scouts expedition. mount gore – see strawberry peak. mount marcel bertrand 72ø (72°55.0´n 25°34.3´w). mountain in se suess land, ne of lumskebugten, the present gravhøjen. the name was used by eugéne wegmann during the 1931–34 tre årseks peditionen (wegmann 1935), and given for marcel alex andre bertrand [1847–1907], a french stratigrapher and structural geologist who made extensive studies in the french jura and alps. mount of gods mercie 69ø (69°03.0´n 26°49.0´w). mountain probably corresponding to the blåserk of the icelandic sagas and the present rigny bjerg on the blosseville kyst (ryder 1892; tornø 1935). it was seen and named by henry hudson during his 1607 voyage, who described it as a “very high mount, like a round castle, which we called the mount of gods mercie” (asher 1860 p. 3). mount mistake 75ø (75°26.4´n 20°58.6´w). minor peak north of ardencaple fjord climbed in error by mike banks and richard brookes in 1952 during the 1952–54 british north greenland expedition; they had to descend it again to regain their route to matterhorn (banks 1955). mount petersberg 76ø (76°09.0´n 18°39.9´w). highest part of the cliff on the south side of trækpasset, store koldewey, named by louise boyd who climbed it on 15 august 1938. it features in den grønlandske lods (1968) as st. petersburg bjerg. mount röhling – see mont röhling. mount shrivenham – see shrivenham. mountain 1 73ø (73°01.8´n 25°20.9´w). informal name used by eha (1953) for a mountain in east suess land, in his report on work during lauge koch’s 1947–49 expedition. mountain 2 73ø (73°02.0´n 25°27.3´w). informal name used by eha (1953) for a mountain in east suess land, in his report on work during lauge koch’s 1947–49 expedition. mountains of the dead – see de dødes bjerg. mountnorris fjord 72ø-8 (72°21.0´n 22°20.0´w; maps 3, 4; fig. 12). fjord on se traill ø. named by william scoresby jr. in 1822 as mountnorris inlet in honour of lord mountnorris. this was possibly lord george annesley, earl of moutnorris [1769– 1844], noted for his voyages to india and ceylon. (mountnorris einbucht, mount norrisfjorden.) mozart dal 76ø-327 (76°35.9´n 23°43.8´w; map 4; fig. 21). long valley in dronning louise land running from farimagsdal to l. bistrup bræ. one of the names given by the 1952–54 british north greenland expedition for composers, it commemorates wolfgang amadeus mozart [1756–91], regarded as one of the greatest musical geniuses. mt. – see mont, monte, mount, mountain. mudderbugt 70ø-57 (70°34.7´n 25°48.8´w; map 4). bay on the se coast of milne land. so named by carl ryder’s 1891–92 expedition because it was very shallow and so full of clay and sand it was impossible to land in their boat. mühldorfer spids 71ø (71°49.7´n 25°24.3´w; map 5). mountain on the south side of the col between spærregletscher and upper bjørnbo gletscher. climbed by karl m. herligkoffer’s 1966 expedition on 19 august, and named after the bavarian town of mühldorf, hometown of edelwald hüttl, one of the climbers. munatius plancus tinde 70ø-383 (70°13.8´n 29°56.0´w). peak 1067 m high on the south side of kaskadesø, west gåseland. it was climbed, and so named, by eduard wenk during lauge koch’s 1958 expedition to honour the founder of the city of basel on its 2000 years anniversary. lucius munatius plancus founded the roman colony of augusta raurica near basel (the present augst) in about 27 bc. wenk was based at the university of basel. münchner tinde 71ø (71°51.3´n 25°23.4´w; map 5). mountain about 2500 m high on the east side of the upper basin of spærre gletscher, stauning alper. climbed by karl m. herligkoffer’s 1966 fig. 60. musk oxen are common in low-lying areas of northern east greenland, where their only enemies are wolves and polar bears. freezing conditions sometimes create a crust of ice on melting snow that the musk ox cannot break through, leading to mass starvation in the affected areas. in this family group the bull is the large musk ox to the right. 258 expedition, and named after münchen (munich), the capital city of bavaria, and home town of karl herligkoffer. mundingshytten 75ø-95 (75°56.0´n 19°56.5´w). danish hunting hut on the south side of the mouth (= munding) of bessel fjord, built by nanok in september 1932. this hut and a norwegian hut nearby (perka hytta) are sometimes referred to as mundingshytter i besselfjorden. now a ruin. (mundinghytten.) munich glacier 72° (72°10.1´n 25°16.2´w; map 5). minor glacier on the south side of vikingebræ. munin 74ø-290 (74°24´n 21°39´w). mountain on the west side of odin dal, th. thomsen land. the name originated from the wintering party at kulhus during the 1931–34 treårseks pedi tionen. see also munin sø. munin sø 71ø-419 (71°07´n 24°21´w). lake in jameson land south of fegin elv. named during the 1967–72 ggu scoresby sund expeditions by svend funder, in style with nearby fegin elv and lodin elv. munin and hugin were odin’s two ravens in old nordic mythology, which every morning left his shoulder, returning to tell him what was happening in the world. munkekutten 70ø-425 (70°39.8´n 28°34.6´w). mountain 1555 m high north of rolige bræ. so named by laurent jemelin during the 1967–72 ggu scoresby sund expeditions because the summit ice cap was reminiscent of a monk’s cowl. munotbjerg 73ø-633 (73°05.9´n 24°52.7´w). mountain about 1150 m high on sw ymer ø, east of margerie dal. named during the 1931–34 treårsekspeditionen by eugène wegmann as munot moun tain, after the castle in schaffhausen, switzerland. murbjerg 74ø-56 (74°28.7´n 19°29.8´w). mountain 853 m high on the north side of dronning augustadalen in wollaston forland. named by karl koldewey’s 1869–70 expedition as mauer berg, possibly because of the steep wall-like appearance of its north side. frebold (1932) used the variation hügel mauern for the same feature. murchison bjerge 72ø-25 (72°17.8´n 25°09.1´w; maps 4, 5). group of mountains in the nw stauning alper. named by a.g. nathorst’s 1899 expedition as murchisons berg after sir roderick impey murchison [1792–1871], a british geologist most noted for ‘the silurian system’ published in 1838. (murchisons bjerge, murchison mountains, murchisonfjella, murchison bjærge.) murgangsdalen 72ø-118 (72°58.4´n 25°55.0´w; map 4). valley in south suess land. named by eugéne wegmann for the mud slide (= murgang) of 1932, caused by the sudden emptying of a glacier-dammed lake (murgangssø) in inner suess land. the slide filled the entire valley with water several metres deep, and washed away wegmann’s camp on 17 august. he was stranded here for four days. an account of the incident is found in koch (1955). (murgang valley, murgangstal.) murgangssø 72ø-451 (72°56.5´n 26°19.3´w; map 4). ice-dammed lake west of østre and vestre spærregletscher in central suess land. this was the lake whose sudden drainage in 1932 washed away eugéne wegmann’s camp (see murgangsdalen). a geological party led by wegmann visited the interior of suess land, and named the lake, in august 1933. (murgangssee.) murray ø [immikkeertikajiit martik] 71ø-10 (71°32.7´n 21°43.2´w; maps 3, 4). island off north liverpool land, named murray island by william scoresby jr. in 1822 after his respected friend admiral murray. (murrays ö, murray-ø, murray insel.) murtinderne 71ø-32 (71°12.0´n 21°47.2´w). mountain in central liverpool land. it was named by william scoresby jr. in 1822 as pinnacle mount, and described as consisting of six or seven tall parallel chimneys forming a beautiful series. the name was adopted in its present form by the place name committee in the 1930s. (pinnacle bjerg.) muschelbjerg 75ø-10 (75°10.6´n 19°51.9´w). range of hills in south hochstetter forland, with two main peaks, nordre and søndre muschelbjerg. named by karl koldewey’s 1869–70 expedition as muschelberg for the finds of fossil shells. the name was preserved by the place name committee in the slightly danicised form muschelbjerg. it is occasionally encountered in the translated but unapproved form muslingebjerg. the cairn at the north end of the ridge of nordre muschelbjerg was said to have been erected by a koldewey party (nyholm-poulsen 1985). (muschelbjærg.) musk ox tower 70ø (70°49.6´n 26°05.3´w). minor summit about 1450 m high on the north side of korridoren, milne land. climbed by the 2004 west lancashire scouts expedition. muskox plateau 70ø (70°29.6´n 22°44.5´w). name used by hermann aldinger (1935) during the 1931–34 treårsekspedi tionen for a plateau area between ostreaelv and lakseelv in south jameson land, and given for the musk ox. muskox pond 76ø (76°13.9´n 18°35.9´w). lake on store koldewey where sampling was undertaken for phytoplankton studies (cre mer et al. 2005). muskox river 70ø (70°27.2´n 22°45.0´w). minor river in south jameson land draining into ostreaelv near its mouth. so named by hermann aldinger (1935) during the 1931–34 treårsekspedi tionen, after the musk ox. (upper muskox river.) muslingebjerg 75ø (75°10.6´n 19°51.9´w). translation into danish used by surlyk (1977) for muschelbjerg in wollaston forland, who also used nordre muslingebjerg and søndre muslingebjerg for the two peaks officially known as nordre and søndre muschelbjerg. see also muschelbjerg. muslingebjerget 76ø (76°07.0´n 18°38.1´w). name occasionally used for aucellabjerget on southern store koldewey during the 1906–08 danmark-ekspeditionen (j. løve, personal communication 2009). muslingeelv 70ø-103a (70°31.5´n 23°02.3´w). river in southern jameson land. named by hermann aldinger during the 1931–34 treårsekspeditionen as mussel river, for the rich finds of fossil lamellibranchs. it was given the name hesteelv by mistake on the 1965 geodætisk institut maps. there appears to be some doubt as to whether the name is officially authorised. muslingeelv 71ø-370 (71°23.4´n 24°36.8´w). small river draining south into nordostbugt, near sydkap. so named by the 1962 oxford university expedition for finds of shells on terraces at the mouth of the river. muslingefjeld 77ø-113 (77°05.9´n 21°42.4´w). mountain in okselandet, north of the west end of sælsøen. named by the 1938–39 mørkefjord expedition, presumably for finds of shells (musling = mussel). it was first visited by paul gelting and alwin pedersen in may 1939. (muslingefjæld.) muslingehjørnet 70ø-352 (70°07.6´n 22°14.5´w). ridge adjacent to bopladsdalen, kap brewster. name used in a report by hassan (1953) describing work on material collected during lauge koch’s 1951 expedition, and given for the numerous fossil shells. myalinadal 71ø-401 (71°34.0´n 22°55.0´w). valley on sw wegener halvø. named by katherina perch-nielsen during the 1967–72 ggu scoresby sund expeditions after the fossil mussel ‘myalina’, common in the valley. myggbukta 73ø-39 (73°29.4´n 21°33.4´w; map 4; fig. 13). norwe gian radio and weather station on the north side of macken zie bugt, an appropriate name as the area is one of the worst for mos quitoes (= mygg) in this part of east greenland. the original sta tion was erected and so named by johan a. olsen in 1922, but the ship carrying his expedition home in 1923, the anni 1, was crushed and lost with all hands on the way through the pack ice. the station was repaired by gunnar isachsen in 1924, next occupied in 1926 by the foldvik expedition, and was entirely rebuilt in 1930. it was manned continuously from 1926 to 1942, and with jónsbú formed part of the norwegian contribution to the inter national polar year 1932–33. in september 1940 the radio equipment was destroyed by the patrol boat fridtjof nansen, and myggbukta was in bad condition at the end of the war. in the summer of 1946 it was re 259 paired, and operated until 1959 when it closed down with the cessation of norwegian state subsidies. the name was approved by the danish authorities in its norwegian form. (myggabuktahuset, mygg-bukta.) myggbukta 73ø (73°27.0´n 21°30.0´w). norwegian hunters name for mackenzie bugt, in use from about 1922 to 1930. see also myggbukta. (midge bay, mosquito bay.) myggedal 73ø-393 (73°32.5´n 25°29.2´w). valley south of grejs dalen in andrée land. named during lauge koch’s 1948–50 expeditions by erdhardt fränkl, but shown only on his cross-section (fränkl 1953). myggesø 72ø-224 (72°10.0´n 23°46.9´w). small lake at hestepas, west of the mouth of mesters vig. named by prospecting teams associated with lauge koch’s 1948–49 expedition. myggvatna 73ø (73°29.0´n 21°42.0´w). swampy area with many small lakes west of myggbukta; so named on an nsiu map (1932a) because it is the breeding ground of mosquitoes. myrvoldhytten 75ø (75°19.9´n 20°18.0´w). norwegian hunting hut built for arktisk næringsdrift on the north side of peters bugt in 1948. it was named after bjarne myrvold, who built the hut together with eigil amsjø. mysteriedalen 73ø-617 (73°15.8´n 28°09.8´w). n–s-trending valley at the west end of knækdalen, named by louise boyd’s 1933 expedition as mystery lakes valley because j.m. wordie’s mystery lakes that he had seen from the summit of petermann bjerg in 1929 were found by boyd to be situated in the valley. mysteriesøer – see øvre mysteriesø, nedre mysteriesø and mystery lakes. mystery lakes 73ø (73°16.1´n 28°08.9´w). two lakes in mysteriedalen on the south side of jættegletscher. j.m. wordie’s 1929 expedition had seen the lakes from the summit of petermann bjerg, in a view over previously unknown land, later explored by louise boyd in 1933. the two lakes are now known as øvre mysteriesø and nedre mysterisø, although the usage mysteriesøer for both lakes is occasionally seen. mythen 70ø (70°14.9´n 29°00.7´w). name used by wenk (1961) throughout his report for lille myteklippe and store myteklippe, two small but distinctive mountains in western gåseland. they were named for their resemblance in shape and geology to grossen mythen and kleinen mythen in canton schwyz, switzerland. mythotinde 72ø (72°12.2´n 25°07.9´w; map 5). peak 2224 m high in the northern stauning alper on the north side of vikingebræ. climbed by claude rey’s 1970 expedition. mæchel-stua 72ø (72°23.1´n 25°15.1´w; fig. 61). norwe gian hunting hut at kap mæchel, at the junction of alpefjord and forsblad fjord. built by the møre expedition in august 1930, and originally known as sentralen, the hut has been regularly repaired and largely retains its original design. it is also known as kap mæchelhytten. märjelen see 73ø (73°44.1´n 27°23.9´w). name used by odell (1937a) for the present madum sø on the north side of gerard de geer gletscher, for a resemblance to the most celebrated of euro pean ice-dammed lakes, the märjelensee held up by the aletsch gletscher in switzerland. möbius bjerg 75ø-5 (75°54.8´n 20°38.6´w; map 4). mountain on the south side of bessel fjord, sw of trums ø. named by karl koldewey’s 1869–70 expedition as cap möbius, after karl august möbius [1825–1908], a german professor of zoology who con tributed one of the zoology sections to koldewey’s narrative. (møbius bjerg.) mønselv 70ø-104 (70°30.0´n 22°53.8´w; map 4). river in south jameson land, draining south to enter the sea west of kap stewart. so named by laurits bruhn during the 1931–34 treårsekspedi tionen after the island of møn, denmark. mønstedhus 75ø-97 (c. 75°42´n 19°33´w). danish hunting station in roseneathbugt, on the north side of langelv delta, built by nanok in 1938 with the aid of funds provided by ‘otto mønsteds fond’. it was manned in the periods 1938–41, 1946–47 and 1951– 52. the station was used for unsuccessful experiments with mink and fox farming, as well as traditional forms of hunting. by the summer of 1953 erosion had removed so much of the coast that the station was in danger, and j.g. jennov with nanok hunters moved it 20 m back from the sea. the station has occasionally been known as danske roseneath to distinguish it from the nearby norwegian station ottostrand, also called norsk roseneath. (mønsted station, mønstedhus station.) mörefjellet 73ø (73°53.8´n 20°05.9´w). sw summit of jackson ø, named after the 1930–32 møre expedition which hunted in this region. the name is found on an nsiu (1932a) map. mörepynten 73ø (73°53.3´n 20°07.3´w). cape on sw jackson ø below mörefjellet. named after the 1930–32 møre expedition which hunted in this region. the name is found on an nsiu (1932a) map. mørepynten 74ø (74°08.7´n 20°28.9´w). small peninsula on the coast of se clavering ø, the present basaltkap. so named on the nsiu maps of lacmann (1937) after the møre and romsdal district of norway, which sent out numerous hunting expeditions to east greenland. (mörepynten.) mørkebjerg 73ø-652 (73°32.6´n 24°55.7´w; map 4). mountain fig. 61. mæchel-stua, the norwegian hunting hut built at kap mæchel, between alpefjord and forsblad fjord, in 1930. 260 1580 m high in east andrée land. so named during the 1931–34 treårsekspeditionen by th. johansen because it is formed of dark (= mørke) rocks. mørkebjerghytten 73ø (c. 73°34´n 24°52´w). norwegian hunting hut in andrée land, ne of mørkebjerg, built for arktisk næringsdrift in september 1933. now disappeared. it was also known as brandalhytten and geologhytten. mørkefinger 72ø-458 (72°50.8´n 28°19.7´w). mountain 2354 m high on the west side of inner agassiz dal. the name was used by eugéne wegmann during the 1931–34 treårsekspeditionen, and was given for its appearance (= dark finger). (markefinger.) mørkefjord 76ø-24 (76°56.4´n 21°09.6´w; map 4). narrow fjord incised into daniel bruun land. named by the 1906–08 danmarkekspeditionen as mørke fjord, because of an unpleasant voyage along the long and narrow fjord in 1906. vigfusdalfjord has been used for the same feature. (mörkefjord, dark fjord, sinus obscurus, mørke fiord, mörke fjord, dimmifjörður.) mørkefjord station 76ø (76°55.7´n 20°19.4´w; map 4). danish scientific station built in 1938 north of the mouth of mørkefjord, west of hvalrosodden. so named by the 1938–39 mørkefjord expedition. it was manned from 1938 to 1941, the last two years because the danish meteorogical institute had requested a continuation of weather reports and because eigil knuth had planned a continuation of expedition activities; the latter was prevented by the outbreak of war. the station is now a ruin. (mørkefjord-station, mørkefjordstation, mørkefjord.) mørkefjordsbugten 76ø-185 (76°56.3´n 20°52.3´w). bay at the entrance to mørkefjord and pustervig. the name was first used by the 1932 gefion expedition. mørkefjordselv 76ø-154 (76°58.3´n 21°41.2´w). river running into the head of mørkefjord, so named by j.p. koch’s 1912–13 expedition. (mørkefjords-elven, mørkefjord-bach.) mørkefjordshytten 76ø-193 (76°56.4´n 20°48.5´w). danish hunting hut on the north side of mørkefjord; it was sailed to this location from hvalrosodden by nanok in august 1933. now a ruin. (mørkefjordhytten.) mørkefjordsplateau 76ø-231 77ø-22a (77°00.0´n 21°19.0´w; map 4). high plateau area between mørkefjord and sælsøen. named by the 1938–39 mørkefjord expedition. (dimmafjarðarhálændi.) mørkeklint 75ø-83 (75°04.8´n 21°05.0´w). cliff in northern th. thomsen land. the name originated from the wintering party at kulhus during the 1931–34 treårsekspeditionen, and was given for its dark colour. mørkholmen 72ø (72°45.3´n 21°57.3´w). small island off the coast of se geographical society ø. used on the nsiu maps of lac mann (1937), the name commemorates rolf mørk [b. 1907], a norwegian artist who took part in the 1933 nsiu expedition to east greenland. møya 71ø (71°45.5´n 25°31.0´w; map 5). mountain about 2350 m high on the sw side of orion gletscher. climbed by the 1996 norwegian stauning alper expedition, and named after a resemblance to a mountain of the same name in northern norway. møysalen 71ø (71°45.3’n 25°29.2´w; map 5). twin-peaked mountain with summits 2450–2500 m high on the sw side of orion gletscher. climbed by the 1996 norwegian stauning alper expedition, and named after a resemblance to a mountain of the same name in northern norway. mågebjerg 71ø-271 (71°56.0´n 23°52.9´w). mountain in the werner bjerge. named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk, and initially published in the forms möwenburg and maagebjerg (bearth 1954, 1959), the name was given for the gulls (= maage, måge). the mountain was climbed by bearth in 1953. mågeelv 70ø-198 (70°31.2´n 21°57.6´w). river in south liverpool land west of mågefjeld draining south into hvalrosbugt. named during the 1931–34 treårsekspeditionen by laurits bruhn. (maageelv.) mågefjeld [qinngajivata qaqqartivaa] 70ø-199 (70°31.6´n 21°54.0´w). hill north of scoresbysund town in south liverpool land. it was named during the 1924–25 colonisation expedition after the numerous gulls (pedersen 1926). (maagefjeld, gullfjeld.) mågefjeldet 76ø-355 (76°44.3´n 21°19.2´w; map 4). coastal cliff in daniel bruun land south of port arthur. the name was recorded by pedersen (1942) who observed here 40 pairs of glaucous gulls in 1938. there was still a glaucous gull colony here in 1989. mågefjeldet 80ø-36 (80°25.9´n 16°13.9´w; map 4). mountain in ne holm land. named by eigil nielsen during the 1938–39 mørkefjord expedition as maagefjeldet, because it is a breedingplace for gulls, and as a counterpoint to mallemukfjeldet in se holm land. (maagefjæld.) mågefjeldet 70ø (c. 70°27´n 26°15´w). name used by helge vedel in his diary of carl ryder’s 1891–92 expedition, for a hill with a gull colony on danmark ø (gulløv 1991). mågegletscher 80ø-37 (80°27.7´n 16°28.2´w; map 4). glacier in ne holm land, west of mågefjeldet. named by eigil nielsen during the 1938–39 mørkefjord expedition in the form maage gletscher. maageholmen 74ø (74°30.0´n 18°57.0´w). name used by the 1908–09 floren expedition for a small island off kap wynn, so named after the many gulls. it was also called lagerholmen. mågensfjeld 81ø (81°18.7´n 14°09.4´w). hill in nw kilen, kron prins christian land, where there are colonies of ivory gull. the name is found on a coloured geological map of kilen printed in 1991 (pedersen 1991). mågenæs 74ø-204 (74°59.8´n 21°45.0´w; map 4). peninsula on the north side of central grandjean fjord. the area was first visited by gunnar seidenfaden in august 1932 during the 1931–34 treårs ekspeditionen, and was named after the colony of gulls (= måge), the only one then known in the region. the name occurs first as a botanical reference locality in gelting (1934) in the form maage næs. balders hage has also been used. mågenæshytten 74ø (74°59.8´n 21°45.0´w). danish hunting hut at the head of the bay at mågenæs, central grandjean fjord. built by nanok in august 1948. mågesøer 76ø-238 (76°48.7´n 19°08.9´w). these are two small lakes on winge kyst, southern germania land, which were named by the 1906–08 danmark-ekspeditionen after the icelandic gull and glaucous gull (måge = gull), both common in the region. (maagesø.) mågetuen [immikkeertaa] 71ø-203 (71°32.7´n 26°11.2´w; map 4). small island on the north side of central nordvestfjord. so named by the 1963 geodætisk institut expedition because it resembled one of the grass-covered mounds (= tuer) which gulls like to perch on. gulls (= måge) also nest on the island. mågeungen 73ø-141 (73°46.0´n 20°24.0´w). small island in carls havn, east hold with hope, so named because of the numerous gulls. the name appears on the nsiu (1932a) map in the form skårungen. mål glacier 71ø (71°55.8´n 24°48.0´w; map 5). name used by the 2007 smc east greenland expedition for a major western branch of storgletscher, on their maps marked as ‘mål glacier’. målebjerg 73ø-644 (73°34.9´n 27°07.5´w; map 4). mountain 1873 m high in western andrée land. so named by ove simonsen during the 1931–34 treårsekspeditionen because it was the surveying station location from which the largest number of points were measured. (maalebjerg.) maanedalpingos 72ø (c. 72°43´n 23°15´w). informal name used by müller during lauge koch’s 1954–55 expeditions, for six pingos he studied in månedal (müller 1959). månedal 72ø-81a (72°42.9´n 23°13.9´w; map 4). valley on north traill ø, south of rold bjerge. so named during the 1931–34 treårsekspeditionen by ove simonsen because of the occurrence 261 of pingos resembling small moon craters. (maanedalen, moon valley.) månegletscher 70ø-259 (70°10.9´n 24°08.9´w). glacier east of soltemplet on volquaart boon kyst, so named during the 1931–34 treårsekspeditionen by laurits bruhn for its association with the nearby names soltemplet and solgletscher (måne = moon, sol = sun). månegletscher 72ø-339 (72°28.0´n 22°07.2´w). small glacier on se traill ø, so named during lauge koch’s 1956–58 expeditions by h.p. heres for its nearly circular shape. månesletten 73ø-408 (73°16.6´n 25°50.6´w). high plain in south andrée land east of junktiondal, so named by erdhardt fränkl during lauge koch’s 1948–50 expeditions for its desolate character, like the surface of the moon. (maanesletten.) månevig 80ø-124 (80°32.5´n 20°30.0´w; map 4). inner e–wtrending part of ingolf fjord, kronprins christian land. named during operation groundhog 1960 together with solvig (vig = bay, måne = moon, sol = sun). måsungane 73ø (73°44.6´n 20°26.0´w). skerry on the south side of carlshavn, so named on an nsiu map (1932a) after the numerous gulls. måtten 76ø-90 (76°41.7´n 18°32.4´w). small island south of danmark havn, south of kap bismarck. named by the 1906–08 danmark-ekspeditionen as maaten. (maaten ø, the mat.) n. polar bear nunatak 69ø (69°12.0´n 32°36.0´w). reference name for a nunatak in the prinsen af wales bjerge, northern kong christian ix land (nielsen et al. 2001). it records an incident when a geologist’s camp was visited by a polar bear. n1, n2, n3, n4, n5, n6, n7 72ø (c. 72°12´n 23°54´w; map 5). designations used on 1:15 000 scale maps of the mesters vig region printed in 1951, for seven rivers west of calamites elv flowing northwards into noret. n naajat 77ø-119 (76°59.6´n 20°21.3´w). cliff se of svingnæs on the west side of southern sælsøen, noted for its gulls and geese. named by the 1938–39 mørkefjord expedition from the greenlandic word for gulls, originally in the form naujat. naasut 74ø-219 (74°01.5´n 21°29.8´w). minor ravine in nw hold with hope, draining river 10. named during the 1931–34 treårsekspeditionen by eigil nielsen as nasutdal, after the grass. (naussut.) nadel klipper 70ø (70°35.5´n 22°38.1´w). name used on a map in wegener (1932) for the present neill klinter, a cliff on the west side of hurry inlet. it probably arises from a mis-reading of ‘neill’. nail glacier 74ø (74°39.1´n 22°28.2´w). tributary glacier to paster ze on its south side. the name was used by battle (1952). nákâkajik – see nakkaakajik. nakkaakajik 70ø-273 (70°05.5´n 23°02.1´w). small glacier on volquaart boon kyst. recorded by the 1955 geodætisk institut name registration, the name means ‘that which falls down’, a reference to active calving of the front. (nákâkajik.) nakkehoved 70ø-238 (70°50.2´n 21°43.3´w; map 4). peninsula on the east coast of liverpool land, north of the mouth of horsens fjord. so named by laurits bruhn during the 1931–34 treårseks peditionen after the headland of the same name in north sjælland, denmark. namsdalsstua 73ø (73°02.6´n 24°42.4´w). norwegian hunting hut built for arktisk næringsdrift in august 1934 at the mouth of fladedal, south ymer ø, by ole klokset and magne råum. the latter was from the namsdalen district of norway. the hut has also been known as flatdalshytta, karl jakobsens bugt and firmanns dalen. (namdalshytten.) nannabreen 74ø (74°15.5´n 20°51.7´w). glacier on central clavering ø. used only on nsiu maps (lacmann 1937), and named after nanna, wife of balder in old nordic mythology. nannut qeqertaat [bjørneøer] 71ø-42 (71°07.0´n 25°25.0´w). island group north of milne land. recorded by the 1955 geo dætisk institut name registration, the name is a translation of the existing danish name, meaning ‘bear islands’. (nánut qeqertait.) nanok 75ø-60 (75°08.5´n 19°44.9´w; map 4). danish hunting station on the south coast of hochstetter forland, built in 1929. the name commemorates the hunting company nanok (nanok = polar bear). a radio station was added in 1931. the station was manned from 1929 to 1941, and intermittently in the period 1946– 55. it was often referred to by hunters as hochstetter, and occasionally as kap rink. (stationen nanok, nanok hunting station.) nanok ø 76ø-171 (76°20.0´n 20°33.3´w; maps 2, 4). island in sw dove bugt. the name was given by the place name committee in 1940, to commemorate the activities of østgrønlandsk fangst kompagni nanok. it was a replacement of the name tuxensø, sug gested by nanok, but rejected by the committee. (nanoks ö.) nanortalik 73ø-536 (73°07.5´n 25°44.9´w). locality at the mouth of nanortalikdal in ne suess land, so named by lauge koch because he killed an unusually large bear here on 15 november 1926, and the next day met three more bears. the greenlandic name translates as ‘the place where there are many bears’. the hut at the mouth of the valley is often known as nanortalik or nanor talikhytten (see bjørneheimen). nanortalik pass 73ø (73°00.4´n 25°47.5´w). name occasionally used by eha (1953) for dalføret, the pass at the crest of nanortalikdal in suess land. nanortalikdal 72ø-139, 73ø-627 (73°01.8´n 25°46.1´w; map 4). valley in suess land extending from north of lumskebugten to nanortalik at the coast of antarctic sund. the name was used by eugéne wegmann during the 1931–34 treårsekspeditionen, and first appeared on maps as nanortalik valley. nanortalikhytten 73ø (73°07.6´n 25°44.4´w). norwegian hunting hut at nanortalik, suess land, built in 1934 for arktisk nærings drift. it was originally known as bjørneheim. nansen-hytten – see nils hermans hytte. nánut qeqertait – see nannut qeqertaat. náparsímavîp nûa – see napparsimmaviip nuaa. náparutikajik – see napparutikajik. napassorssuaq [kirkespiret] 74ø-40 (74°41.2´n 18°31.6´w). moun tain 497 m high on lille pendulum with a spire-like summit. the name is essentially a translation of the danish name, meaning ‘the upright-standing’. naportoqs elv 70ø-122 (70°55.3´n 22°37.5´w). river at the head of hurry inlet, ne of eli bjerg. named by alfred rosenkrantz during lauge koch’s 1926–27 expeditions as naportok river, after his greenlandic assistant, eli napartok. nappangulikajiip kangersiva 70ø (70°26.8´n 21°48.8´w). name recorded by the scoresbysund local newspaper in 1984 for hartz vig, also known as kangertivatsiaakajik, the bay between kap tobin and kap swainson. nappangulikajik 70ø (70°26´n 21°45´w). name recorded by the scoresbysund local newspaper in 1984 for the kap swainson area, and apparently also for the point known as napparutikajik. napparsimmaviip nuaa 70ø-369 (70°29.0´n 21°57.3´w). cape on the east side of scoresbysund town. recorded by the 1955 geo dætisk institut name registration, the name translates as ‘hospital cape’. it is due south of the old hospital building, which was originally the french expedition house of 1932–33. (náparsímavîp nûa.) napparutikajik 70ø-334 (70°25.8´n 21°44.5´w). point on the coast a little west of kap swainson, south liverpool land. recorded by the 1955 geodætisk institut name registration, the name means ‘the little upstanding’, and refers to a cairn. nappangulikajik was reported by the scoresbysund local newspaper in 1984 as in use for this feature, and the general area of kap swainson. (náparuti ka jik.) 262 narhvalgletscher 72ø-171 (72°46.8´n 25°18.4´w). large glacier on the lyell land side of narhvalsund, named by louise boyd’s 1937 expedition as narhval glacier. detailed studies were made here by richard foster flint. narhvalsund 72ø-41 (72°46.4´n 25°06.4´w; maps 3, 4). sound between ella ø and lyell land. named by a.g. nathorst’s 1899 expedition as narhvalssundet because they were surprised to see a flock of narwhales here (fig. 8). today they are a not uncommon sight in many east greenland fjords. (narwhal strait, narwhal sound, narwhale sound, narhval sund.) narrow ridge 73ø (73°30.9´n 23°20.8´w). locality between two ravines on the south side of sederholm bjerg, gauss halvø. the name was used in a report on work during the 1931–34 treårseks pe ditionen (johansson 1935). (smala ryggen.) narsakajik 70ø (70°27.6´n 22°22,8´w). name recorded by tuborg & sandell (1999) for a locality about 1 km west of the settlement kap hope / ittaajimmiit that is the site of inuit ruins. the name trans lates as ‘the little plain’. nathorst bjerg 73ø-113 (73°23.9´n 23°04.0´w). mountain on the sw coast of gauss halvø. it was named during the 1931–34 treårs eks peditionen by gunnar säve-söderbergh as mt. nathorst after alfred gabriel nathorst [1850–1921], who led an expedition to east greenland in 1899 to search for andrée’s lost balloon expedition (see andrée land). nathorst discovered and mapped much of the fjord complex between latitudes 72° and 74°n, and made a number of notable geological observations. norwegian maps of the 1930s used ketilfjellet for the same feature. nathorst elv 70ø (70°48.7´n 22°42.1´w). name occasionally used for the river in the n–s-trending valley west of nathorst fjeld, on the west side of hurry inlet (e.g. lilliesköld & salvigsen 1991). nathorst fjeld 70ø-130 (70°49.5´n 22°39.6´w). mountain on the west side of hurry inlet, west of the fame øer. named by g.c. amdrup’s 1898–1900 expedition after a.g. nathorst (fig. 62), whose 1899 expedition was the first to reach the head of hurry in let. see also nathorst bjerg. (nathorst mountain, mount nathorst, mont nathorst.) nathorst fjord 71ø-48 (71°41.0´n 22°28.5´w; maps 3, 4). fjord between canning land and wegener halvø, discovered by g.c. amdrup’s 1898–1900 expedition and named after a.g. nathorst (fig. 62). see also nathorst bjerg. (nathorsts fjord, nathorst fiord, nathorst fjorden.) nathorst gletscher 73ø-714 (73°08.1´n 28°16.6´w). glacier between nathorst tinde and mona bjerg, western frænkel land. the name was first used in climbing and geological reports of louise boyd’s 1933 expedition (odell 1934a, 1937a, 1939), and approved in 1952 following explorations in the region by john haller and eduard wenk. see also nathorst bjerg. nathorst land 71ø-145 72ø-80a (71°50.0´n 26°30.0´w; maps 3, 4). extensive land area bounded to the east by alpefjord, prinsesse gletscher and borgbjerg gletscher, to the north by forsblad fjord and tærskeldal, and to the south by inner nordvestfjord and f. graae gletscher. named by lauge koch during the 1931–34 tre års ekspeditionen. koch had mapped the area during reconnaissance flights in 1932. see also nathorst bjerg. nathorst tinde 73ø-539 (73°06.9´n 28°18.0´w; fig. 65). mountain 2372 m high west of nordenskiöld glet scher, named by j.m. wordie in 1929 as nathorst peak after a.g. nathorst. nathorst had mistaken this peak for petermann bjerg in 1899 (wordie 1927). the first ascent was made by neill odell and walter wood during louise boyd’s 1933 expedition. see also nathorst bjerg. (nathorst bjærg.) nathorst valley 70ø (70°48.7´n 22°42.1´w). name occasionally used for the n–s-trending valley west of nathorst fjeld, on the west side of hurry inlet (lilliesköld & salvigsen 1991). nattvika 72ø (72°57.8´n 24°50.9´w). bay on east maria ø. the name is used as a reference locality in nsiu botanical and zoological reports (vaage 1932). naujat – see naajat. naussut – see naasut. navnløs 73ø-719 74ø-203 (74°00.0´n 22°13.9´w). river in ne hudson land, named during the 1931–34 treårsekspeditionen by gunnar säve-söderbergh as unnamed river. nebalopokrygge – see nipiluttut. nebbøyra 72ø (72°56.3´n 21°58.5´w). narrow penisula in east geographical society ø, nw of kap mackenzie. so named on the nsiu maps of lacmann (1937) for the beak-shaped form (nebb = beak). nedre antarctic gletscher 71ø-250 (71°57.9´n 23°49.4´w; map 5). glacier in the eastern werner bjerge, flowing from antarctic pas ne into the head of kolledal. named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk (nedre = lower). nedre arkosedal 71ø-303 (71°35.8´n 24°45.0´w; map 5). valley draining ne to the front of bjørnbo gletscher, with deep red arkosic sandstone on both sides. named by enrico kempter during lauge koch’s 1956–58 expeditions. nedre funddal 72ø-191 (72°06.8´n 24°06.1´w; map 5). valley in north scoresby land, draining ne into mesters vig. named by prospecting teams associated with lauge koch’s 1948–49 expeditions for finds of lead ore. ‘funddal’ has occasionally been used as a common name for both nordre funddal and nedre funddal (nordre = northern, nedre = lower). nedre gefionelv 72ø-186 (72°09.5´n 24°09.6´w; map 5). river in north scoresby land on the ne side of schéele bjerg, which joins øvre (= upper) gefionelv just before reaching store blydal. named by prospecting teams associated with lauge koch’s 1948–49 expeditions after gefion, goddess of nordic mythology, who changed her four sons into oxen and ploughed out the danish island of sjælland from sweden. nedre mysteriesø 73ø-616 (73°16.0´n 28°08.0´w). lower of two lakes in mysteriedalen. in 1933 louise boyd distinguished j.m. wordie’s mystery lakes as upper mystery lake and lower mystery lake. nedre randgletscher 71ø-287 (71°52.1´n 24°11.2´w; map 5). western and lower of two glaciers south of aldebaran gletscher, on the north flank of randspids. named during lauge koch’s 1953– 54 expeditions by peter bearth and eduard wenk. nedre rypegletscher 73ø-545 (73°01.4´n 28°11.5´w). lower part of rypegletscher, north goodenough land, named by j.m. wordie in 1929 as lower ptarmigan glacier. nedre studer gletscher 71ø-244 72ø-307a (72°00.2´n 23°51.2´w). glacier in the north werner bjerge. named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk after bernhard rudolf studer [1794–1887]. he was a swiss mathematician and structural geologist, a pioneer of alpine geology who be came professor at bern university, and is noted for stimulating the first geological mapping of switzerland. negeren 75ø-53 (75°10.5´n 19°58.3´w). mountain in south hoch stetter forland, on the north side of søndre muschelbjerg. the name originated from the wintering party at kulhus during the 1931–34 treårsekspeditionen, and was given for the colour of the coal outcrops (negeren = the negro). idahöhe has also been used. negritaelv 77ø-62 (77°28.8´n 20°58.7´w). stream draining into the head of v. clausen fjord, inner skærfjorden. so named during the 1931–34 treårsekspeditionen by david malmquist after the brand of rum (negrita) they drank during the surveying expedition. neild bugt 71ø-7 (71°21.9´n 21°50.2´w; map 4). small fjord or bay in liverpool land. it was named neild bay by william scoresby jr. in 1822, probably, like other features in the northern part of liverpool land, after manchester friends. neill klinter 70ø-138 (70°35.5´n 22°38.1´w). range of cliffs on the west side of hurry inlet. named by william scoresby jr. in 1822 as 263 neill’s cliffs, after patrick neill [1776–1851], a naturalist who became head of the large printing firm of neill & co., which printed scoresby’s two-volume work on the arctic regions for archibald constable. scoresby describes the cliffs as 300 feet high, and appears to have intended the name to apply to the cliffs just north of kap stewart (fig. 3). the name is now used in a wider sense for the cliffs extending between kap stewart and constable pynt which are up to 500 m high. (neill cliffs, neill’s klipper, neill falsen, nadel klipper.) nell sø 73ø-587 (73°58.8´n 24°16.4´w). lake in south ole rømer land, named by sigurd skaun and harald welde in 1932 as nells vann. girl’s name. neptune glacier 71ø (71°38.5´n 25°30.1´w). glacier in the south stau ning alper, the present løberen, which drains south into nordvestfjord. named by james clarkson’s 1961 expedition after the planet neptune, eighth major planet from the sun. this name is in common use in mountaineering literature. nerdiit iaat 70ø (70°31.6´n 21°54.0´w). name recorded by the scoresbysund newspaper in 1984 as in local use for mågefjeld, the hill north of the town, also officially known as qinngajivata qaqqartivaa. it translates as ‘the place of the geese’. nerlerit inaat [constable pynt] (70°44.3´n 22°38.2´w). greenlan dic name for the location of the airfield at constable pynt built in 1985 to serve the oil and gas exploration on jameson land. the name derives from the original greenlandic designation nerterit inaat kangittiit recorded in 1955. (nertiit inaat.) nerterit inaat kangittiit 70ø-159 (70°44.3´n 22°38.2´w). delta of ugleelv at the head of hurry inlet. one of the names recorded by the 1955 geodætisk institut name registration, the name translates as ‘the inner dwelling place of the wild geese’. (nerterit inait kangig tît.) nerterit inaat kitteq 70ø-131 (70°45.7´n 22°38.7´w). broad delta where gåseelv enters the west side of hurry inlet. the name was recorded by the 1955 geodætisk institut name registration, and translates as ‘the outer dwelling place of the wild geese’. (nerterit inât kíteq.) nerterit inait kangigtît – see nerterit inaat kangittiit. nerterit inât kíteq – see nerterit inaat kitteq. nertivit kangersivat [gåsefjord] 70ø-17 (70°10.0´n 27°15.0´w; maps 3, 4). large e–w-trending fjord between gåseland and milne land. the greenlandic name appears in this form on modern maps but was formerly oqqummut kangertiva. nes-odden 74ø (74°12.1´n 21°53.1´w). norwegian hunting hut on sw clavering ø, built by the foldvik expedition in 1927. it was replaced by a new hut 200 m to the west in 1954 known as storholts hus. the hut has also been known as øtkerhytten or kap øtker. nesodden 72ø (72°48.0´n 22°07.1´w). peninsula on east geograph ical society ø on the south side of inner cambridge bugt. used only on nsiu maps (lacmann 1937), and named for the shape (nes = peninsula). it is a common norwegian place name. nesvatnet 72ø (72°45.9´n 21°59.9´w). lake behind the peninsula dragneset on se geographical society ø. so named on lacman’s (1937) maps. neue hütte – see hansa bugt. neuhausendal 71ø (71°50.8´n 23°18.2´w). valley on the north side of ørsted dal, apparently the present horsedal. so named during the 1936–38 two-year expedition by hans stauber (1940), after the swiss town of neuhausen near schaffhausen, the home of his assistant hans hübscher. neviatiakdal – see niviarsiaq. nevis 71ø (71°39.5´n 25°20.3´w; map 5). mountain about 2150 m high at the head of jupiter gletscher, south stauning alper. it was first climbed by james clarkson’s 1961 expedition, and probably named after ben nevis, the highest mountain in scotland. new mountains 68ø (69°00.0´n 29°30.0´w). name used by gino watkins for the present watkins bjerge, situated almost entirely south of latitude 69°n. the mountain range was observed during a flight along the coast in 1930. see also watkins bjerge. new valley 72ø (72°53.8´n 27°33.4´w). name used by bretz (1935) for the present bocksrietdalen in his geology report of louise boyd’s 1933 expedition. louise boyd explored and mapped the valley in 1931 and 1933. newnham pas 71ø-368 (71°56.5´n 25°16.5´w; map 5). pass about 2350 m high between the head of cantabræ and newnham glacier, stauning alper. named by the 1963 cambridge expedition after newnham college, cambridge, established in 1875 as the second women’s college. (newnham col.) newnham glacier 71ø (71°54.3´n 25°15.5´w). glacier in the central stauning alper, south of newnham tump, so named by the 1963 cambridge east greenland expedition. it was later named ravnas bre by a norwegian expedition. newnham tump 71ø (71°55.8´n 25°14.8´w; map 5). peak 2500 m high on the roslin gletscher – cantabræ divide, sw of newnham pas. climbed and named by the 1963 university of cambridge expedition. newton klippe 77ø-130 (77°00.3´n 24°52.8´w; map 4). prominent cliff on the north side of admiralty gletscher, nw dronning louise land. one of the names given by the 1952–54 british north greenland expedition for notable scientists, it commemorates the british physicist and mathematician sir isaac newton [1642– 1727]. he is considered the culminating figure of the 17th century scientific revolution, and among numerous achievements is noted for the three fundamental laws of mechanics and invention of the infinitesimal calculas. nid-bjerge 72ø (72°02.4´n 24°08.4´w). name used by styger (1951) for a ridge in the werner bjerge, north of vestre gletscher, in his report on a climbing excursion during lauge koch’s 1950 expedition. fig. 62. alfred gabriel nathorst [1850–1921], was a swedish paleobotanist and geologist who took part in five expeditions to spitsbergen and greenland. his 1899 expedition to northern east greenland in search of traces of s.a. andrée’s lost balloon expedition mapped an extensive region between latitudes 72° and 74°n. 264 nidelv 72ø-296 (72°04.0´n 24°05.2´w; map 5). minor river in the werner bjerge draining east into deltadal, which rises on the north side of the ridge named nid-bjerge by styger (1951). the name was used during lauge koch’s 1948–50 expeditions by peter bearth and eduard wenk. nidsdal 72ø (72°04.0´n 24°05.2´w). name used by pessl (1962) for the valley in which nidelv flows. niels hansen næs 75ø-55 (75°08.5´n 19°53.0´w; map 4). peninsula just west of nanok hunting station. the name came into use in the 1930s by danish hunters, and was given for niels hansen [1878–1963], known usually as ‘gamle niels’ or ‘niels ivigtut’. he was employed at ivigtut for nine years, worked as a carpenter during the establishment of scoresbysund in 1924–25, and from 1925 until 1940 hunted with nanok. he was a member of the sledge patrol from 1940 to 1945. the locality has sometimes been referred to as niels hansens næse, or simply næsen (= the nose). (nils hansens næs.) nielsnæs 74ø (74°09.1´n 20°25.7´w). cape on the coast of se clavering ø, east of basaltkap. the name occurs on a sketch map in gustav thostrup’s 1921 logbook, and may have been given for the first mate on the dagny in 1921, niels larsen sleth. niesen 74ø-355 (74°39.1´n 20°30.4´w). mountain in nw wolla ston forland. so named during lauge koch’s 1936–38 expeditions by wolf maync and andreas vischer because it resembled in shape and geology the mountain of the same name south of spiez in the berner oberland, switzerland. (mt. niesen.) niflheim 75ø-82 (75°25.2´n 21°32.8´w; map 4). mountain in northern c.h. ostenfeld land south of smallefjord. the name originated from the wintering party at kulhus during the 1931–34 treårsekspeditionen, and was given because surveying here was greatly delayed by fog. niflheim was a world of mists in old nordic mythology, which existed before the earth was created. niggli dal 73ø-632 (73°13.4´n 26°40.7´w). valley in east frænkel land south of niggli spids. so named by eugéne wegmann during the 1931–34 treårsekspeditionen because the valley was very dull in appearance, and paul niggli was said to be a dull lecturer (f. schwarzenbach, personal communication 1996). paul niggli [1888–1953], a swiss petrologist for many years professor at the mineralogisch-petrographische institut zurich, was noted for his scale of hardness. see also niggli spids. niggli spids 73ø-628 (73°15.7´n 26°40.4´w; map 4). mountain in east frænkel land. so named by eugéne wegmann during the 1931–34 treårsekspeditionen after paul niggli. niggli was known for his belief in magmatic rather than migmatitic processes, and wegmann is said to have given the name intentionally so that he could hold lectures with the title ‘the migmatites of niggli spids’ (f. schwarzenbach, personal communication 1996). see also nig gli dal. niinngarpik [grøfteelv] 70ø-185 (70°31.2´n 22°23.5´w). river in south liverpool land draining into hurry inlet. recorded by the 1955 geodætisk institut name registration, the name translates as ‘the place where something bends’, and is explained by an incident when a man wading the river appeared to have bent legs due to refraction. (nîngarpik.) niklausdal 72ø-443 (72°39.7´n 27°23.8´w). valley in gletscher land on the nw side of skræntdal. name used by eugéne weg mann during the 1931–34 treårsekspeditionen. see also buri søer. nils hermans hytta 72ø (72°53.9´n 24°22.7´w). norwegian hunting hut on the north side of vega sund, se of svedenborg bjerg. built for arktisk næringsdrift in august 1929 by sverre sørensen and thor halle, it was named after halle’s son nils herman. it has also been known as lindqvist-hytta and nansen-hytta. nils holgersen nunatakker 73ø-717 (73°23.8´n 29°52.0´w). nuna tak group west of martin knudsen nunatakker. named by john haller following explorations on a catalina flight during lauge koch’s 1953 expedition. the name derives from a childrens’ story by the noted swedish writer selma lagerlöf, translated into danish as ‘niels holgersens vidunderlige rejse gennem sverige’. the hero of the original swedish version is nils holgersson, and the name of the nunataks is a mixture of danish and swedish. nîngarpik – see niinngarpik. nioghalvfjerdsbræ 79ø (79°33.0´n 21°00.0´w). name often used for the glacier filling nioghalvfjerdsfjorden, between lambert land and hovgaard ø (e.g. weidick 1995). the glacier is afloat, and the large lake blåsø on the north side of the glacier is tidal. see also nioghalvfjerdsfjorden. nioghalvfjerdsfjorden 79ø-3 (79°33.0´n 21°00.0´w; maps 1, 4). fjord between lambert land and hovgaard ø entirely filled by floating glacier ice. so named by the 1906–08 danmark-ekspedi tionen because it lies at latitude 79°n. the name was originally regarded as temporary, but acquired a new significance in the diaries of jørgen brønlund as the last presumed resting place of mylius-erichsen and høegh-hagen, such that it was necessary to keep it. the bodies of the two men were long thought to have been left on the glacier ice, or on one of the small islands at the front of the glacier, but despite a series of search expeditions in recent years neither their bodies nor their lost diaries have been found. (niog halv fjerds-fjord, seventy-nine fjord, 79-fjord, nioghalvfjerds fiord.) nioghalvtredskilometernæsset – see fyrretyvekilometernæsset. niogtredivekilometernæsset – see fyrretyvekilometernæsset. nipilugtut – see nipiluttut. nipiluttut 74ø-218 (74°00.5´n 21°30.3´w). series of minor ridges in nw hold with hope, at the head of rivers 11 and 12. they were named during the 1931–34 treårsekspeditionen by eigil nielsen as nebalopokrygge, ‘because the mountains howl’. (nipilugtut.) nippoldt gletscher 74ø-160 (74°02.5´n 22°29.6´w). small glacier in the nørlund alper draining north into wordie bugt, named by lauge koch’s 1929–30 expeditions. lacmann’s (1937) maps use a. schmidtbreen for this glacier. (nippoldts gletscher.) nissedal 70ø-194 (70°35.5´n 22°03.7´w). small valley draining into jættedal, south liverpool land. so named by laurits bruhn during the 1931–34 treårsekspeditionen for its relatively small size (nisse = pixie). nisseelv 70ø-195 (70°35.5´n 22°03.7´w). river in nissedal, south liverpool land, so named during the 1931–34 treårsekspedi tionen by laurits bruhn. niviarsiaq 74ø-217 (74°00.3´n 21°26.5´w). minor ravine in nw hold with hope, in which river 13 flows. named during the 1931–34 treårsekspeditionen by eigil nielsen, originally as neviatiakdal. probably named after the willow herb, the national flower of greenland, which in greenlandic is niviarsiaq (= the virgin or maiden). niviarsiat 73ø-537 (73°04.0´n 25°13.7´w). mountain on the south side of antarctic sund, named during lauge koch’s 1926–27 expeditions (koch 1929a). the mountain is formed by strongly folded exposures of the eleonore bay group, and the name derives from the vivid colours, some of which are reminiscent of the willow herb – see niviarsiaq. (niviarssiat, niviarsiak.) nivlheimdalen 74ø (74°24.2´n 20°57.0´w). valley on north clav ering ø, the present skilledal. so named on the nsiu maps of lacmann (1937) after the niflheim of old nordic mythology, a world of mists which existed before the earth was created. no-name-dal 71ø (71°53.8´n 22°53.1´w). name used by university of dundee expeditions in the 1970s for the valley west of regn buedal draining into fleming fjord. noa dal 73ø-623 (73°19.4´n 25°03.2´w). valley on ymer ø between noa sø and dusén fjord. named after noa sø, the name came into general use during the 1931–34 treårsekspeditionen. noa pas 73ø (73°19.5´n 25°15.0´w). low pass between blomster bugten and noa sø, west ymer ø. the name was used by eha (1953) during work on lauge koch’s 1947–49 expeditions. noa sø 73ø-569 (73°19.3´n 25°10.7´w; map 4; figs 35, 74). lake on 265 west ymer ø, between blomster bugten and dusén fjord. named noa lake during lauge koch’s 1929–30 expeditions by gunnar seidenfaden and arne noe-nygaard, after the danish natural history society ‘noa’ (natur historiske onsdags aftener). noaelven 73ø (73°19.4´n 25°03.2´w). name used by andersen (1937) and others for the minor river in noa dal draining eastwards from noa sø to dusén fjord. noahytten 73ø (73°19.1´n 25°02.8´w). name sometimes used for the norwegian hut at the mouth of the river draining noa sø, at the head of dusén fjord. it was built in august 1932 by the crew of the isbjørn for salmon fishing, and is also known as bunnhuset, holm boe-hytta and laksehytten. nok 74ø-395 (74°24.8´n 24°21.7´w; map 4). isolated mountain or nunatak 1555 m high in eastern bartholin land. named by john haller following explorations during lauge koch’s 1956–58 expeditions, after an austrian word for a mountain massif. noorajik 70ø-291 (70°27.5´n 22°16.9´w). cape east of ittaajimmit [kap hope], sw liverpool land. the name was recorded by the 1955 geodætisk institut name registration, and translates as ‘the little cape’. (nôrajik) noorajik kangitteq [kap hope] 70ø-286 (70°27.7´n 22°22.9´w). cape in sw liverpool land. recorded by the 1955 geodætisk institut name registration, it means ‘the western little cape’. (nôrajik kangigteq.) noorajiva 70ø-319 (70°26.4´n 21°58.4´w). point on the east coast of rosenvinge bugt. one of the names recorded by the 1955 geo dætisk institut name registration, it translates as ‘its little cape’. the scoresbysund newspaper recorded in 1984 the usage noorn gaviva kangideq for this feature. (nôrajiva.) noorajiva 71ø-209 (71°18.1´n 25°08.1´w). peninsula west of sydkap at the mouth of nordvestfjord. the name was recorded by the 1955 geodætisk institut name registration, and means ‘its little peninsula’. (nôrajiva.) noorngaviva kangideq 70ø (70°26.4´n 21°58.4´w). name recorded by the scoresbysund newspaper in 1984 for noorajiva, a point on the east coast of rosenvinge bugt. it translates as ‘the westernmost cape’. noorngaviva kiddeq 70ø (70°26.0´n 21°58.2´w). name recorded by the scoresbysund newspaper in 1984 for nuugaatsiaq kitteq, a point on the east coast of rosenvinge bugt. it translates as ‘the easternmost cape’. nôrajik, nôrajik kangigteq – see noorajik, noorajik kangitteq. nôrajiva – see noorajiva. nord gletscher 71ø (71°55.5´n 23°55.8´w). name occasionally used by bearth (1959 p. 21) for a glacier in the werner bjerge, possibly the present hvidefirn (nord = north). nordborgen 72ø (72°44.5´n 24°27.5´w). norwegian hunting hut built in september 1935 by the suløya expedition at the north foot of kongeborgen, traill ø. it was the northernmost hut in their hunting district, with a roof formed by an upturned boat. (nord borg hytten, nordborghuset, norborg.) nordbugten [immikkeertaata kangertiva] 71ø-36 (71°35.0´n 26°27.2´w; map 4). short fjord or large bay on the north side of central nordvestfjord. named by carl ryder’s 1891–92 expedition as nordbugt. nordelv 72ø-103 (72°38.8´n 25°13.6´w). river in the northern half of polhem dal draining north into narhvalsund. named by ove simonsen during the 1931–34 treårsekspeditionen. nordelv 77ø-112 (77°08.6´n 20°41.4´w; map 4). river flowing north into the south end of annekssøen, named by the 1938–39 mørkefjord expedition. it is close to sydelv that flows south into sælsøen. nordenskiöld bjerg 71ø-24 (71°36.3´n 22°33.4´w). mountain in canning land named during lauge koch’s 1926–27 expeditions as mt. nordenskiöld. the name was clearly intended to commemorate the work carried out in the vicinity by otto nordenskjöld (fig. 63), whose name koch consistently spells ‘nordenskiöld’ (koch 1929a); the letters ‘i’ and ‘j’ were interchangeable in old danish. noe-nygaard (1934) used the correct spelling ‘nordenskjöld’ for both the mountain and the geological formation named after the mountain, but the original mis-spelling is now firmly established in the literature. nils otto gustaf nordenskjöld [1869–1928], a swedish explorer, geologist and oceanographer, was professor of geology at the university of gothenburg from 1905. he led expeditions to greenland in 1900 and 1905, and was leader of the 1901–03 swedish antarctic expedition. nordenskiöld bugt 75ø-26 (75°14.1´n 18°04.4´w; map 4). bay on the east coast of shannon. named by karl koldewey’s 1869–70 expedition as nordenskjöld bucht, probably after nils adolf erik nordenskiöld [1832–1901], the noted swedish arctic explorer (fig. 64). see also nordenskiöld gletscher. the alternative (or mis-spelling) ‘nordenskjöld’ occurs on all koldewey’s maps, and is also found in contemporary german biographical works (e.g. pog gen dorff 1863). (nordenskjølds bugt, nordenskiøld bay.) nordenskiöld gletscher 73ø-524 (73°02.1´n 28°25.6´w; maps 3, 4; fig. 65). major glacier at the head of kejser franz joseph fjord, named by a.g. nathorst’s 1899 expedition after niels adolf erik nordenskiöld [1832–1901]. nordenskiöld (fig. 64) had encouraged nathorst to take up his perhaps most noted work on the fossil flora of skåne, and nathorst also took part in nordenskiöld’s 1883 expedition to west greenland. nordenskiöld was most noted for the first voyage through the ne passage and around asia in the vega. (norden skjölds gletscher, nordenskiöld glacier, nordenskiöld breen.) nordenskiöld ø 72ø-69 (72°39.7´n 22°28.9´w; map 4; fig. 14). island at the mouth of vega sund, named by a.g. nathorst’s 1899 expedition as kap nordenskiöld. white (1927) interpreted nat horst’s cape as an island which he renamed nordenskiold island, although it is possible nathorst may have intended the name to apply to the present kap mcclintock 17 km east of the present island. the name is not directly attributed to n.a.e. norden skiöld, and it is possible that nathorst had intended to honour his son gustaf erik nordenskiöld [1868–95], an archaeologist and mineralogist; nathorst had described collections of fossils made by g.e. nordenskiöld in spitsbergen in 1880 (higgins 1986). (norden skjölds ø, nordenskiöldøya.) nordfjord 73ø-511 (73°42.0´n 24°17.0´w; maps 2–4). n–s-trending fjord 13 km wide between strindberg land and gauss halvø. named nordfjorden by a.g. nathorst’s 1899 expedition for its direction. a hut on the east side of the fjord is sometimes known as nordfjord (see brehytta). (north fjord, north fiord, nordfiord.) nordfjordhuset 73ø (73°42.1´n 24°30.6´w). scientific station built in 1931 on the east coast of strindberg land during the 1931–34 treårsekspeditionen. it is sited immediately south of strindberg hytten. nordhoek bjerg 73ø-54 (73°47.3´n 22°06.5´w; map 4). mountain 1502 m high on the west side of loch fyne, named by h.g. backlund during lauge koch’s 1929 expedition in the form mt. nordhoek after the captain of the expedition ship godthaab, hannes gysbert nordhoek [1894–1953]. nordhoek was first mate on the godthaab in 1924 during its search for the teddy, captain of the godthaab during koch’s expeditions in 1929 and 1931, captain of the sværdfisken from 1932 to 1939, and in post-war years until 1952 was captain of the ships disko and umanak mainly serving west greenland towns. (nordhoekberg.) nordkap 78ø-37 (78°54.1´n 19°16.1´w; map 4). northern cape of schnauder ø in jøkelbugten. named by the 1938–39 mørkefjord expedition. nordkjosen 72ø (72°44.8´n 22°01.0´w). bay on se geographical society ø, south of cambridge bugt. used on the nsiu maps of lacmann (1937), the name derives from a place of the same name in the troms district of norway. 266 nordlige christian den ix’s land – see kong christian ix land. nordlige frederik den viii’s land – see kong frederik den viii land. nordlige fligely hytten – see fligelyhytten. nordlige jægersundhytte 76ø (76°19.0´n 20°48.3´w). norwegian hun ting hut built on the island tvillingerne, sw dove bugt, in august 1933. jægersund is the channel between tvillingerne and nanok ø. the hut has also been called kroken, vestre tvillingen and tvillinghytten. nordmarken 77ø-102 (77°45.0´n 21°00.0´w; maps 1, 2, 4). exten sive land area between kofoed-hansen bræ and skærfjorden, north of søndermarken. named by the 1938–39 mørkefjord expedition, who considered it the northern segment of an expanded germania land. nordneset 72ø (72°03.4´n 23°06.3´w). name used by the 1930–32 møre expedition for a peninsula near the hunting station at ant arc tic havn, possibly the present kap syenit (rogne 1981). nordostbugt [kangerterajiva] 71ø-33 (71°20.0´n 24°41.0´w). shallow bay east of sydkap, usually filled by stranded icebergs. named by carl ryder’s 1891–92 expedition as nordöst bugt because of the general trend. (nordøst fjord, nordostbugten, nord öst bugten, north-east bay, north-east fjord, northeast bay.) nordostrundingen 81ø-71 (81°21.2´n 11°30.3´w; maps 1, 4). point on the ne coast of kronprins christian land, named by the 1906–08 danmark-ekspeditionen. this is the easternmost point of greenland, and for the sledge parties of the 1906–08 danmarkekspeditionen travelling northwards was the point where the coast began to curve westwards. nordprofil 74ø (74°44.4´n 20°00.1´w). geological reference locality on se kuhn ø, used by maync (1947) in his description of work during lauge koch’s 1936–38 expeditions. nordre basisdal 71ø-102 (71°38.3´n 22°16.9´w). valley in se canning land draining north to ålborg fjord. the name appears to have first been used by säve-söderbergh (1937) in the form n. basis valley, and derives from surveying work during lauge koch’s 1936–38 expeditions. nordre biot – see biot-stua. nordre fligelyhytten – see fligelyhytten. nordre funddal 72ø-190 (72°07.6´n 24°09.5´w; map 5). valley in north scoresby land, a north branch of nedre funddal. named by prospecting teams associated with lauge koch’s 1948–49 expeditions for minor finds of lead ore in quartz veins. nordre gneisnæs 76ø-158 (76°16.3´n 18°34.4´w; map 4). nor thern of two gneiss ridges bounding areas of sediments on the east side of store koldewey. used by the 1906–08 danmark-ekspedi tionen in the form northern gneiss naze. (northern gneissnæs, nordliche gneisnaes.) nordre koldewey ø 76ø (76°39.0´n 18°40.9´w). name sometimes used during the 1906–08 danmark-ekspeditionen for lille kolde wey, which is situated to the north of store koldewey (e.g. amdrup 1913). (north koldewey island.) nordre muschelbjerg 75ø-51 (75°10.9´n 19°48.6´w). slightly more northern part of muschelbjerg, situated ene of søndre muschelbjerg, hochstetter forland. so named by hans frebold during the 1931–34 treårsekspeditionen. an unapproved danicised version of the name, nordre muslingebjerg, was used by surlyk (1977). fig. 63. nils otto gustaf nordenskjöld [1869–1928] was a swedish geologist, geographer, and polar explorer. he was particularly noted for his leadership of the 1901–04 swedish antarctic expedition, aboard the ship antarctic, and also led expeditions to greenland in 1900 and 1905. nordenskiöld bjerg on canning land was named after nordenskjöld, but the misspelling used by lauge koch is preserved in the name. fig. 64. nils adolf erik nordenskiöld [1832–1901], the noted swedish arctic explorer, was a geologist, mineralogist and geographer. he was most noted for the first successful voyage through the north-east passage in the vega in 1878–79. in northern east greenland, nordenskiöld bugt and nordenskiöld gletscher were both named after n.a.e. nordenskiöld. 267 nordre muslingebjerg – see nordre muschelbjerg. nordredepot ø 78ø-17 (78°12.8´n 20°29.0´w; map 4). island in jøkelbugten, variously referred to in the 1906–08 danmark-eks pedi tionen reports as nordre depot and nordre depot island. the northern of two depots was placed here in october 1906. nordsylen 72ø-247 (72°20.5´n 24°33.1´w; map 5). northernmost spire of the syltoppene, north stauning alper. the name was given by the place name committee as a substitute for birgitsbjærg, a name proposed by erdhardt fränkl during lauge koch’s 1950–51 expedition. nordvestelv 70ø-33 (70°44.4´n 25°26.3´w). tributary to nordøst elv on east milne land, south of charcot havn, named during the 1931–34 treårsekspeditionen by hermann aldinger as nordwest fluss. nordvestfjord [kangertertivarmiit kangertivat] 71ø-37 (71°15´n 25°10´w to 72°15´n 28°30´w; maps 3, 4). very long fjord extending nw from the north end of hall bredning. so named by carl ryder’s 1891–92 expedition because of its direction (fig. 7). the distance from the mouth of scoresby sund via hall bredning to the innermost point of nord vest fjord is 313 km, a continuous stretch of water credited with being the longest fjord in the world. (north-west fjord, nordvest fjorden, nordvest fjord, north west fjord.) nordvestklint 79ø-35 (79°23.2´n 21°25.6´w; maps 1, 4). cliff in nw lambert land. the name is a modification by the place name committee of a 1960 proposal by john haller. nordvestkæret 74ø (74°28.8´n 20°36.1´w). reference locality used by scientists visiting zackenberg forskningsstation. nordvestre havnenæs 76ø-256 (76°45.8´n 18°43.0´w). peninsula on the west side of danmark havn. it appears on the 1906–08 danmark-ekspeditionen maps in the form nv. havnenæs (e.g. johansen 1912). nordøstelv 70ø-36 (70°45.0´n 25°21.9´w). river north of kap leslie, east milne land, draining into charcot bugt. named by hermann aldinger during the 1931–34 treårsekspeditionen as nordost fluss. nordøstgrønlands nationalpark the north-east greenland na tio nal park, established in 1974 and expanded westwards across north greenland in 1988, is the largest national park in the world with an area of 972 000 km2. the southern boundary extends from latitude 71°n north-eastwards and north along the east margin of the stauning alper to 72°n. most of the park comprises a large part of the inland ice, but the coastal regions include the main breeding areas of the musk ox and polar bear. nordøstplateau 70ø-35 (70°45.4´n 25°22.7´w). plateau north of kap leslie, east milne land, overlooking charcot bugt. named during the 1931–34 treårsekspeditionen by hermann aldinger as nordost plateau. nordøstre havnenæs 76ø-257 (76°45.8´n 18°39.2´w). peninsula on the east side of danmark havn, originally denoted in the form nö. havnenæs on maps of the 1906–08 danmark-ekspeditionen (e.g. johansen 1912). norejva 69ø (69°54.6´n 22°58.7´w). name used by sølberg (1980) for a peninsula on the south side of steward ø, on the west side of a bay where two houses were built in 1971–72 (norejva = noorajiva = its little cape). noret 72ø-86 (72°13.3´n 23°52.1´w; maps 4, 5; fig. 66). lagoon with a narrow entrance on the south side of kong oscar fjord near mesters vig. named by ove simonsen during the 1931–34 treårs ekspeditionen. it is a common danish place name for an enclosed bay or lagoon. noret pools 72ø (c. 72°13´n 23°47´w). name used by the university of dundee expeditions between 1968 and 1974 for six small lakes near mesters vig, on the east side of noret. norma hytta 74ø (74°09.2´n 21°30.4´w). norwegian hunting hunt probably built by normann andersen in 1953–54 at svampebugt, sw clavering ø. it was named after andresen’s daughter norma. norma-øien 71ø (c. 71°45´n 23°36´w). name used by ingstad (1935) for a small hill rising from a flat valley floor, possibly in ørsted dal or pingel dal. it was named after the daughter of his companion normann andersen. (norma island.) norske villa – see villaen. norske roseneath – see ottostrand. lille petermann petermann bjerg nathorst tinde mona bjerg gog magog nordenskiöld gletscher kejser franz joseph fjord kjerulf fjord knækdalen fig. 65. nordenskiöld gletscher at the inner end of kejser franz joseph fjord, with the entrance of kjerulf fjord visible in the left foreground and knækdalen at the right. on the skyline a series of high summits are visible: from left, lille petermann (2709 m), petermann bjerg (2971 m), nathorst tinde (2382 m), mona bjerg (c. 2300 m high), gog (2628 m) and magog (2521 m). 268 norske petersbugt station – see jónsbú. norske øer 78ø-14 79ø-26 (79°04.0´n 17°50.0´w; maps 1, 4). one large and several small islands east of the front of zachariae isstrøm, so named by the 1906–08 danmark-ekspeditionen. j.p. koch (1916) records that the name is to be considered a compliment to the two norwegian members of the expedition, h.l. hagerup and k.j. ring. (norske öer, norwegian islands, norske islands, norske öarne.) norskepashytten 74ø (74°25.6´n 20°20.9´w). norwegian hunting hut built by the foldvik expedition in 1927 on the ne side of zackenberg bugt, wollaston forland. it was originally known as gisvold. it was given this name to distinguish it from the nearby danish hut known as pashuset. norskeryggen 71ø-390 (71°26.9´n 23°17.6´w). ridge in jameson land east of olympen, of which the highest point is pelion. the name was suggested by russel marris, following his journeys in 1968, as he thought it was a former norwegian hunting area. norsketinden 72ø-265 (72°08.1´n 25°03.3´w; maps 4, 5; figs 27, 67). mountain peak 2870 m high in the north stauning alper between vikingebræ and gullygletscher, the second highest peak in the region. it was climbed by the danish–norwegian expedition on 7 august 1954, and originally referred to as erik rødes tinde or eirik raudes tinde. the place name committee proposed the present name as a compromise and a counterpoint to nearby dansketinden. the second ascent was made by wolfgang diehl and fritz schwarzenbach, also in 1954, and the third ascent in 1968. north bay 75ø (75°20.8´n 18°15.8´w). name occasionally used by ejnar mikkelsen for sengstacke bugt, a bay on the north side of shannon, in his report on the 1909–12 alabama expedition (e. mikkelsen 1922). north cirque glacier 73ø (73°33.5´n 27°26.6´w). name used in a report by odell (1937a) for the north tributary of louise gletscher in louise a. boyd land, studied during louise boyd’s 1933 expedition. north gletscher 72ø (72°06.8´n 28°42.3´w). name used on 1951 usaf aeronautical charts for the present f. graae gletscher at the innermost end of nordvestfjord. north-west gletscher has also been used. north lochan 72ø (c. 72°15´n 23°55´w). name used by university of dundee expeditions between 1968 and 1974 for a small pool near langdyssen at the ne end of mestersvig airfield. north river 72ø (72°30.5´n 23°58.9´w). name used by university of dundee expeditions between 1968 and 1974 for a minor stream west of karupelv draining into holm bugt, sw traill ø. north-west gletscher 72ø (72°06.8´n 28°42.3´w). name used on 1957 ams maps for the present f. graae gletscher at the head of nordvestfjord. north gletscher has also been used. northern fault valley – see fault valley. notting hill 72ø (72°08.3´n 24°51.2´w). mountain 2400 m high on the south side of dunottar gletscher, north stauning alper. first climbed by the 1963 imperial college expedition, and named after the london district of north kensington, now best known for its caribbean carnival. there has been confusion over the relative positions of the peaks kensington and notting hill (watson 1964; bennet 1972). nûa – see nuaa. nuaa [kap swainson] 70ø-335 (70°25.9´n 21°43.6´w). cape in south liverpool land. the name was recorded by the 1955 geodætisk institut name registration, and means ‘the cape’. (nûa.) nuclei 74ø (c. 74°12´n 20°49´w). name used by mittelholzer (1941) for the three peaks monacleus, binucleus and trinucleus on clav ering ø, in his report on work during lauge koch’s 1938–39 expeditions. nucleisee 74ø (c. 74°14´n 20°37´w). small lake in grønnedal, east clavering ø, east of the mountain group which mittelholzer (1941) called nuclei. nûgâtsâ – see nuugaatsaa. nûgâtsiaq kíteq – see nuugaatsiaq kitteq. nûgatsiâjik – see nuugatsiaajik. núkaitsoq – see nukkaatsoq. nûkajît akorngáne kangerterajik – see nuukajiit akornganni kanger terajik. nukkaatsoq 70ø-186 (70°32.0´n 22°14.5´w). hill in south liver pool land west of scoresbysund. recorded by the 1955 geodætisk institut name registration, the name translates as ‘that which is believed to have great strength’. (núkaitsoq.) nuldal 72ø-221 (72°07.8´n 23°53.5´w). valley sw of ekspeditions huset draining into mesters vig. named by prospecting teams associated with lauge koch’s 1948–49 expeditions. on the 1:15 000 scale maps of the mesters vig region it is situated between two groups of rivers referred to informally as 1ø–7ø and 1v–8v. (nul dalen.) nulog 73ø (73°16.7´n 24°48.1´w). name used by eha (1953) for the isolated hill on the south side of inner dusén fjord known as rumpen. this was apparently eha’s original suggestion and has the same meaning in greenlandic (nulog = rumpen = the rump). nummer 1 hytten 75ø (75°20.1´n 20°11.9´w). danish hunting hut built in august 1930 by nanok on the north side of peters bugt, and officially known as petersbugthytten. it has also been known as bundhytten. numsen 74ø (74°09.7´n 20°13.9´w). small peninsula on east clav ering ø between kap mary and dahls skær. the name appears on a sketch map in gustav thostrup’s 1921 logbook, and was given for the shape of the peninsula (numsen = backside, bottom). nunataami elv 80ø-61 (80°45.0´n 20°19.0´w). river draining romer sø, which flows through vandredalen to the north inner arm of ingolf fjord. named by elmar drastrup’s 1938–39 expedition as nunatâme elv. drastrup (1945) observed that the name was derived from an inuit dialect word from the kap york district meaning ‘new land’, so that the name translates as ‘the river in the new land’. 1982 nunatak 69ø (69°03.0´n 32°46.0´w). reference name for a nunatak in the prinsen af wales bjerge, northern kong christian ix land (nielsen et al. 2001). geological work was carried out here in 1982. nunatak godfrey 69ø (c. 69°10´n 31°28´w). peak 2585 m high in the lindbergh fjelde, west of christian iv gletscher, northern chri stian ix land. climbed by the 2001 lanchester greenland expedihovedethovedethovedet rypesøen noret aaronip sarpaa fig. 66. view of the lagoon noret, looking west, with the small lake rypesøen at the east end of mestersvig airstrip (not visible in photo). the channel aaronip sarpaa and peninsula hovedet are in the foreground. the john haller photograph collection, geus archive. 269 tion and named after dan godfrey, surveyor on martin lindsay’s 1934 expedition. a surveying spike on the summit was initially thought to have been placed by martin lindsay’s expedition, but this spike marks a fixed point established by the geodætisk in stitut in 1984 or 1986. the expedition altitude measured by gps was 2655 m, about 70 m too high compared to the gi determination. nunatakgletscher 73ø-518 74ø-240b (73°57.4´n 26°00.0´w; map 4). glacier at the head of geologfjord. discovered by a.g. nat horst’s 1899 expedition, and named nunatak glacieren because of the several mountain tops or nunataks which appeared to project from it. (nunatak glacier.) nunatakken 75ø (75°19.1´n 17°47.9´w). rocky prominence forming the east point of kap sussi, which was used as a lookout post by the 1943–44 operation bassgeiger. the name is reported by olsen (1965). it has a small stone wall enclosure on the summit, and is still (1988) connected by a telephone wire to the base camp site. nunatâmeporten 80ø (80°35.0´n 19°10.5´w). mountain 1593 m high on the south side of ingolf fjord, west of brede spærre gletscher. so named by elmar drastrup’s 1938–39 expedition (dra strup 1945) because it was situated at the entrance to the newly discovered inner part of ingolf fjord (see also nunataami elv). nurven 74ø (c. 74°07´n 20°46´w). skerry off the coast of se clav ering ø. the name is used on an nsiu map (1932a). nuua [kap swainson] 70ø-335 (70°25.9´n 21°43.6´w). cape in sw liverpool land. nuugaatsaa [albuen] 70ø-144 (70°34.4´n 22°34.7´w). cape on the west side of hurry inlet. one of the names recorded by the 1955 geodætisk institut name registration, it translates as ‘the rather large cape’. (nûgâtsâ). nuugaatsiaq kitteq 70ø-320 (70°26.0´n 21°58.2´w). point on the east coast of rosenvinge bugt. the name was recorded by the 1955 geodætisk institut name registration, and means ‘the eastern cape’. (nûgâtsiaq.) nuugatsiaajik 70ø-256 (70°26.6´n 23°11.0´w). gravel and sand delta forming a minor cape on the south coast of jameson land. recorded by the 1955 geodætisk institut name registration, the name means ‘the rather large bad cape’. (nûgatsiâjik.) nuukajiit akornganni kangerterajik [gabet] 70ø-218 (70°40.4´n 21°38.8´w). bay on the east coast of south liverpool land, west of rathbone ø, between the capes snuden and hagen. the name was recorded by the 1955 geodætisk institut name registration, and translates roughly as ‘the bay with two bad capes’. (nûkajît akorn gáne kangerterajik.) nuungajiva 70ø (70°24.6´n 21°56.7´w). name for a cape near kap tobin, also known as vardepynten, recorded by the scoresbysund newspaper in 1984. ny jónsbu 75ø (75°14.8´n 20°52.6´w). norwegian hunting station built in 1948 for arktisk næringsdrift on the south side of arden caple fjord as a replacement for the nearby jónsbú station burnt down in 1943. ny jónsbú was manned only from 1948 to 1950, but was maintaind for many years by sirius. it was restored in 1995 by nanok. see also jónsbú. ny mønstedhus 75ø (75°42.1´n 19°33.8´w). hut built in 2002 on the east coast of hochstetter forland from the remains of mønstedhus that had been destroyed by coastal erosion (p.s. mikkelsen 2008). ny station – see dødemandsbugten. ny store snenæs hytten 76ø (76°49.2´n 19°21.3´w). hut built at snenæs on the south coast of germania land in 1999, from prefabricated sections made at danmarkshavn weather station. ny valdemarshaab 74ø (74°18.4´n 20°13.6´w). danish hunting station built in 1923 on the north side of young sund by øst grønlandske fangstkompagni as a replacement for the station valdemarshaab at kap borlase warren – the station is now known as sandodden. the original name was given for a.l. valdemar manniche [1867–1957] (see also valdemarshaab). nyboder 74ø (c. 74°13´n 20°14´w). this name appears on a sketch map in gustav thostrup’s 1921 logbook at the mouth of the present henningselv, and may have been intended for the site of inuit ruins. the hunting hut at this locality was built in 1930 (see henningelvhytten). the name may commemorate the rows of houses of the same name in copenhagen built by christian iv for families of the danish navy. nyhavn 72ø-259 (72°15.5´n 23°55.7´w; maps 4, 5). harbour and bay 2 km north of the airfield at mestersvig, initially used for the landing of goods for the lead mine, and shipping out of ore. the name came into use in about 1950, and first appeared on the detailed topographic maps of the mesters vig region. it had also been used in newspapers reporting the mining activities. the harbour also served the airfield until its closure in 1985. nordisk mineselskab moved a number of barracks from the mining town (minebyen) in store blydal down to nyhavn in 1971, and used nyhavn as a base for prospecting exploration. nymfegryde 72ø-386 (72°02.9.1´n 23°21.5´w). basin-shaped valley on the east side of majdal, north scoresby land. so named by hans kapp during the 1957–58 lauge koch expeditions, for its suggestive nymph-like shape. nyt ekspeditionshus 72ø (72°07.9´n 23°51.7´w). house built on the west side of mesters vig in 1974 or 1975 as a replacement for ekspeditionshus, which was destroyed by an avalanche in the spring of 1973. this new house was originally an office shed at mestersvig airfield, damaged by an aeroplane crash in 1973. (nyt ekspeditions hus.) næsen 80ø-46 (80°31.5´n 20°14.5´w; map 4). cape at the head of ingolf fjord, between solvig and månevig. named by eigil nielsen during the 1938–39 mørkefjord expedition, for its appearance (næsen = the nose). nøglefjeldet 80ø-43 (80°34.9´n 21°00.5´w). mountain on the fig. 67. the highest summits in the stauning alper viewed from the north-east. dansketinden is 2842 m high and norsketinden 2797 m high. the john haller photograph collection, geus archive. norsketinden dansketinden 270 north side of the mouth of sødalen. named by eigil nielsen during the 1938–39 mørkefjord expedition in the form nøglefjældet. it was examined in detail, and regarded as the key (= nøgle) to the geological relationships. (nöglefjældet.) nøkkedal 70ø-171 (70°40.0´n 22°19.0´w; map 4). valley in liverpool land draining west into hurry inlet. named during the 1931–34 treårsekspeditionen by laurits bruhn (nøkke = water elf ). nøkkedal river 70ø (70°39.3´n 22°25.2´w). name occasionally en countered in reports of work during the 1931–34 treårsekspedi tionen for the river draining nøkkedal (e.g. kranck 1935). nøkkefossen 73ø-651 (73°36.7´n 25°08.9´w). river in east andrée land with two waterfalls, draining into geologfjord. named during the 1931–34 treårsekspeditionen by th. johansen. nørlund alper 73ø-56 74ø-21a (74°00.0´n 22°31.8´w; map 4). area of pronounced alpine topography in north hudson land (fig. 15). named by lauge koch’s 1929–30 expeditions in the form nørlund alps after n.e. nørlund [1885–1980]. nørlund was pro fessor at the university of copenhagen from 1922 to 1956, and director of the geodætisk institut from 1923 to 1955. he had been a member of the 1931–34 treårsekspeditionen committee. at his own request the name was not to be officially recognised until after his death. lacmann’s (1937) map used nörlundtindane in a more restricted sense, for the ridge north of rungsted gletscher. (nør lund alpen, nörlund-tindene, nörlundalpen.) nørlund land 75ø-42 (75°42.5´n 21°30.0´w). land area between ardencaple fjord and bredefjord in the south, and bessel fjord to the north. this was one of the new names on the 1932 edition of the geodætisk institut 1:1 million scale map, given by lauge koch following aerial observations during the 1931–34 treårsekspedi tionen. it commemorates n.e. nørlund, then director of the geo dætisk institut (see also nørlund alper), and although officially approved nørlund refused to allow the name to be printed on the institute’s maps during his lifetime. this land area is now part of dronning margrethe ii land. (nørlunds land.) nørre biland 78ø-32 (78°37.0´n 21°48.0´w; maps 1, 4). northern part of hertugen af orléans land, west of nørreland. named by the 1938–39 mørkefjord expedition. (nörre biland.) nørre mellemland 78ø-33 (78°22.6´n 21°12.9´w; maps 1, 4). northernmost but one part of hertugen af orléans land, between blæsebræ and gammel hellerup gletscher. named by the 1938–39 mørkefjord expedition. nørre orienteringsø 76ø-236 (76°49.8´n 19°36.6´w). name used by the 1906–08 danmark-ekspeditionen for the northernmost island of the orienteringsøerne. (nr orienteringsö, northern orien terings island.) nørrefjord 71ø-132 (71°05.6´n 22°07.1´w). fjord on the north side of storefjord, central liverpool land, named during the 1931– 34 treårsekspeditionen by laurits bruhn. nørreland 78ø-31 (78°42.3´n 21°17.5´w; maps 1, 4). northern most part of hertugen af orléans land. named during the 1938– 39 mørkefjord expedition. nørresund 76ø-213 (76°30.7´n 20°56.3´w). sound on the north side of godfred hansen ø. named during the 1938–39 mørke fjord expedition. nørresundbyhytten 76ø-203 (76°33.9´n 20°45.4´w). danish hunting hut on the se coast of andreas lundager ø. it was built by nanok in 1938, and named after the town of nørresundby, near ålborg in denmark. the newspaper ‘aalborg stiftstidende’ had raised funds to support the nanok expeditions. (nørresundby hytten, nr sundby hytten.) nørretop 74ø-400 (74°03.8´n 25°35.0´w; map 4). mountain in north strindberg land, named by hans r. katz during lauge koch’s 1948–49 expeditions. nørvehytta 73ø (73°13.9´n 23°27.6´w). norwegian hunting hut on the north side of dusén fjord, west of kap graah. it was built by arktisk næringsdrift in october 1929, and named after the ålesund merchant elias nörve, a director of arktisk næringsdrift. now a ruin. (nörve, nørve, nørvehytten.) nålene 74ø-59 (74°25.1´n 19°41.8´w). mountain 1142 m high in wollaston forland, named by karl koldewey’s 1869–70 expedition as die nadeln presumably because of its double-spired, pointed summit, although possibly also after an alpine mountain of similar name. (naalene, mt die nadeln.) nålepuden 70ø-266 (70°01.2´n 23°35.4´w). mountain 1713 m high on volquaart boon kyst, so named during the 1931–34 treårseks peditionen by laurits bruhn for its spiked basalt pinnacles (nålepuden = pin cushion). o o. lenz fjelde 77ø-144 (77°11.7´n 20°14.3´w; map 4). part of valdemarsmuren, søndermarken. named by john haller following explorations during lauge koch’s 1956–58 expeditions, after oskar lenz [1848–1925], an austrian geographer and geologist, who had written up geological observations on karl koldewey’s 1869–70 expedition with franz toula. obélix 71ø (c. 71°56´n 25°46´w). prominent granite tower on the east side of prinsessegletscher. named and illustrated in the report on the 1968 claude rey expedition (georges & rey 1969), alt hough it was apparently not climbed. obrutschew bjerg 73ø-287 (73°20.7´n 22°45.1´w). mountain on the sw coast of gauss halvø. the name was used by gunnar säve-söderbergh during the 1931–34 treårsekspeditionen, and commemorates the prominent russian zoologist and vertebrate palaeontologist, dmitri obruchev [1900–1970], an authority on devonian fishes and stratigraphy. (mt. obrutschew.) observatoriehalvø 74ø-48 (74°32.0´n 18°50.2´w). peninsula on south sabine ø, on the sw side of germania havn. named stern wartenhalbinsel by karl koldewey’s 1869–70 expedition because it was the site of edward sabine’s 1823 observatory where he conducted his pendulum experiments. koldewey had searched in vain for the observatory site, and first discovered its location on rereading sabine’s account after the return of the expedition. (observatory peninsula.) odd arnesenfjellet 74ø (74°21.3´n 20°43.4´w). mountain ridge 1238 m high on ne clavering ø, part of the present koralbjerg. so named on the nsiu maps of lacmann (1937) after odd arnesen [1897–1946], a norwegian journalist who worked for the oslo ‘aftenposten’ for 25 years. he was especially interested in the arctic, and edited ‘polar-årboken’ up to 1945. odin dal 74ø-288 (74°53.4´n 21°32.5´w; map 4). valley extending sse from central grandjean fjord across th. thomsen land to svejstrup dal. the name is attributed to the overwintering party at kulhus in 1935, and was given for odin, greatest and most important of all the gods of old nordic mythology. (odins dal.) odinsborg 77ø-60 (77°20.6´n 20°24.9´w; map 4). mountain in ne søndermarken on the south side of c.f. mourier fjord. so named by david malmquist during the 1931–34 treårsekspeditionen. see also odin dal. odinshanesø 70ø-433 (70°34.0´n 27°57.9´w). small lake on sw milne land. named during the 1967–72 ggu scoresby sund expe ditions by max fumasoli after the numerous red-necked phalarope (= odinshane). okse river 72ø (72°04.8´n 23°48.8´w). name used by pessl (1962) for the river in oksedal, se of mestersvig. oksebakkerne 74ø (74°28.6´n 20°27.8´w). low hills ne of zacken berg forskningsstation, where musk oxen often graze. the name is used as a reference locality by visiting scientists. oksedal 72ø-226 (72°04.8´n 23°48.8´w; maps 4, 5). valley se of mestersvig. named by prospecting teams associated with lauge koch’s 1948–49 expeditions after the musk ox. 271 okseelv 74ø (74°28.1´n 20°24.0´w). river in kuhnpasset, wollaston forland, draining sw to zackenberg bugt. the name is used as a reference locality by visiting scientists from zackenberg forsk nings station. (okseelven.) oksefaldet 77ø-116 (77°05.9´n 21°03.1´w). steep cliff on the north coast of sælsøen. so named by the 1938–39 mørkefjord expedition, because paul gelting and alwin pedersen found two dead musk oxen at the foot of the cliff on 10 may 1939. they had fallen down the cliff since their previous visit to the locality. oksehorn 72ø-305 (72°01.5´n 23°39.9´w). mountain between the head of oksedal and kolledal. named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk. okselandet 77ø-107 (77°13.0´n 21°25.0´w; maps 2, 4). triangular land area west of annekssøen, bounded on the west by stor strømmen and to the south by sælsøen. named during the 1938–39 mørkefjord expedition, probably by paul gelting who visited the area in june 1939 and considered it the best area in the region for musk oxen. oksepas 72ø-375 (72°01.7´n 23°43.0´w). pass between the head of oksedal and rødedal, north scoresby land. so named by hans kapp during lauge koch’s 1957–58 expeditions, because musk ox cross the pass here. oksesletten 74ø (74°13.0´n 20°19.0´w). area between henningselv and grønnedal on east clavering ø. the name appears on a sketch map in gustav thostrup’s 1921 logbook, and was given for the musk ox. oksestenen 76ø-297 (76°54.6´n 20°10.2´w). large stone on a cape east of mørkefjord station. so named by the 1938–39 mørkefjord expedition, because musk ox used it as a scratching stone. oksetrappen 73ø (73°45.0´n 20°35.4´w). name used by gelting (1934) for a locality about 3 km west of carlshavn, hold with hope, where a series of marine terraces form a staircase-like fea ture where musk oxen graze. oktoberø 77ø-95 (77°52.4´n 19°06.0´w; map 4). island ne of gamma ø, the northernmost point reached on depot-laying journeys by the 1938–39 mørkefjord expedition in october 1938. ole rømer land 73ø-91 74ø-138 (73°11.5´n 24°20.3´w; maps 2, 4). land area bounded by promenadedal, waltershausen gletscher and vibeke gletscher. on some early maps this general area is part of steno land. the name first appears on the 1932 1:1 million scale geodætisk institut map prepared on the basis of 1932 aerial observations by lauge koch during the 1931–34 treårsekspeditionen. it was named after the noted danish astronomer ole christensen rømer [1644–1710]. (ole rømers land). oleryggen 71ø (71°50.4´n 25°36.8´w; map 5). peak in the ne part of the borgbjerg gletscher region, southern stauning alper. probably first climbed and named by the 1977 schwäbische stauning alper expedition. olestua 76ø (76°07.3´n 19°44.8´w). norwegian hunting station on the east coast of hochstetter forland at kap carl ritter, the most northerly station erected by john giæver’s expedition in 1932. named after ole sivertsen, who helped build the station and manned it with john johnsen from 1932 to 1934. it was accidently burnt down in 1982. carl ritterhytta has also been used, as well as the names beurmann and ullestuen. olgas ø 76ø (76°27.0´n 20°54.5´w). name used by c.s. poulsen during the 1906–08 danmark-ekspeditionen for the present godfred hansen ø in dove bugt (lundbye 1984). òllumlengri 70ø (c. 70°17´n 23°00´w). fjord possibly identifiable with the present scoresby sund (tornøe 1944). the name is mentioned in the icelandic sagas, and means the ‘fjord longer than all other fjords’. while the description in the account of ivar bårdsson admirably fits scoresby sund, other commentators have placed the fjord farther south. the name appears on several old maps against the legendary sound supposed to cut across green land from the west to east coasts, e.g. hans egede’s 1818 and 1846 maps (fig. frontispiece; egede 1818, trap 1928). (allumlengri, ol lum længri fiord, øllum lengre.) olrik pynt 76ø-82 (76°39.2´n 18°38.5´w). minor cape on the east side of lille koldewey. named by the 1906–08 danmark-eks peditionen as olriks pynt, possibly after ejnar olrik of the royal dockyard (j. løve, personal communication 2009). (olrik odde.) olsen creek 74ø (74°02.2´n 21°35.2´w). error for foldvik kløft, found in koch (1929a p. 115) in a reference to the ‘olsen creek formation’ which should have been the foldvik creek formation. foldvik kløft is about 6 km east of kap stosch. the norwegian olsen brothers were hunters based at the krogness station 2 km sw of kap stosch, and had shown lauge koch some of the excellent fossiliferous sections for which the region is now noted. koch at one time appears to have intended to name both the river and the geological formation after the olsen brothers rather than after nils foldvik (svend bendix-almgreen, personal communication 1997). olsen nunatakker 76ø-152 (c. 76°48.6´n 26°30.6´w; map 4). two small nunataks in west dronning louise land. named by the 1909–12 alabama expedition as olsen’s nunatakker after hans p. olsen, one of the members of the expedition who sledged to nw dronning louise land with wilhelm laub in april 1910. oluf kolsrudfjellet 72ø (72°57.1´n 23°23.9´w). mountain ridge on central geographical society ø. used only on nsiu maps (lac mann 1937), and named after oluf kolsrud [b. 1885], a norwegian historian who was professor at oslo university and had interests in developments in greenland. olympelven 71ø-193 (71°18.0´n 23°46.4´w; map 4). river in jameson land draining se from olympen. named during lauge koch’s 1936–38 expeditions by hans stauber. olympen 71ø-183 (71°26.5´n 23°31.1´w; maps 3, 4). high mountain in jameson land, with a summit ice cap. so named during lauge koch’s 1936–38 expeditions by hans stauber after the olym pus of the greek gods. stauber climbed the mountain in august 1938. oqaluppiup ataa 70ø-367 (70°29.0´n 21°58.5´w). coastal stretch south of the church in scoresbysund. one of the names recorded by the 1955 geodætisk institut name registration, it means ‘that lying below the church’. (oqaluugpiup atâ.) oqaluugpiup atâ – see oqaluppiup ataa. oqqummut kangertiva [gåsefjord] 70ø-17 (70°10.0´n 27°15.0´w). large fjord south of gåseland. this is the early name recorded by the 1955 geodætisk institut name registration, and translates as ‘the sheltered fjord’. on modern maps the greenlandic name has been changed to nertivit kangersivat. (orqungmut kangertiva.) orange crest 73ø (73°07.6´n 28°18.7´w). name used in a climbing report (odell 1934a) for a yellowish granite forming the nw ridge of nathorst tinde. ordnungheia 72ø (72°53.4´n 22°19.3´w). mountain on east geographical society ø. used on the nsiu maps of lacmann (1937), and named after franz ordnung [b. 1886], who worked on preparation of the detailed nsiu map sheets of east greenland at hansa luftbild gesellschaft. orelfjellet 74ø (74°19.7´n 21°08.3´w). mountain on central clavering ø. so named on the nsiu maps of lacmann (1937) after eduard von orel [1877–1941], an austrian officer who developed stereophotogrammetric instruments for mapping. (v. orelfj.) orgelpiberne 76ø-14 (76°17.4´n 20°23.7´w; map 4). mountain 740 m high on nanok ø, west of roon bugt. so named by the 1906–08 danmark-ekspeditionen, possibly for the appearance of massive vertical ribs resembling organ pipes (j. løve, personal communication 2009), and perhaps for the incident recorded by friis (1925), who noted that while camped in the vicinity they heard strange, deep tones coming from the cliff with intervals of a few minutes caused by falling rocks. (organ pipes, orgelpiben.) orienteringsnunatak 73ø-409 (73°57.4´n 29°19.4´w). nunatak south of hobbs land. named during lauge koch’s 1951 expedi272 tion by hans r. katz, who climbed it to reconnoitre his route through the nunataks. (rekognoszierungs-nunatak.) orienteringsspids 74ø-62 (74°28.5´n 20°47.4´w). mountain 1342 m high on the north side of tyrolerfjord, se of zackenberg. named by karl koldewey’s 1869–70 expedition as orienterungs spitze, probably because it was a prominent peak used as a surveying point. (orienteringstoppen, mt orienterungsspitze.) orienteringsøerne 76ø-7 (76°47.0´n 19°46.0´w; map 4). island group in dove bugt, one of which was named orienterungs insel by karl koldewey’s 1869–70 expedition, probably because the view from the summit was useful in determining the route of the expedition. the term east island is used in the english edition of kolde wey’s narrative, probably an error of translation. the 1906–08 danmark-ekspeditionen extended the usage of the name to three large and several small islands. (orienterings island, recognition islands.) orion gletscher 71ø-327 (71°44.9´n 25°23.4´w; map 5). glacier in the south stauning alper, flowing se to join jupiter gletscher. named orion glacier by john hunt’s 1960 expedition, after the major constellation. orion-borgbjerg col 71ø (71°47.3´n 25°30.3´w). col between the head of orion gletscher and borgbjerg gletscher. the name is used by bennet (1972). the 1996 norwegian stauning alper expedition crossed the col during their south to north ski traverse, and called it an dorus mor (the great gate). orleans island 77ø (77°50.0´n 18°49.0´w). name occasionally used for the present gamma ø in orléans sund in reports on the 1909– 12 alabama expedition (e. mikkelsen 1922). see also orléans sund. orléans sund 77ø-7 (77°48.0´n 20°00.0´w; maps 2, 4). sound between gamma ø and stormlandet. named by the 1905 duke of orléans expedition as fiord d’orléans. see hertugen af orléans land. orqungmut kangertiva – see oqqummut kangertiva. ortlerspids 74ø-72 (74°22.0´n 21°11.1´w). mountain 1513 m high on north clavering ø, named during karl koldewey’s 1869–70 expedition as ortler spitze by julius payer, because of its resemblance to mountains he had explored in the ortler alps of the austrian tyrol. according to seidenfaden (1931) there is some uncertainty as to the relative positions of this mountain and højnålen. (ortlerfjellet, mt. ortler.) orvaelv 73ø-128 (73°47.6´n 20°41.3´w). river in home forland draining south into tobias dal. named on an nsiu map (1932a) as orva, possibly after a river of the same name in the hedmark district of norway. orvin fjæld 74ø (73°59.9´n 21°30.5´w). name used by eigil nielsen (1935) in a report on work carried out on the 1931–34 treårseks pedi tionen, for part of the mountain west of blåelv, north hold with hope. orvinhytta 73ø (73°05.2´n 23°19.9´w). norwegian hunting hut on the north side of sofia sund, sw of celcius bjerg. built in sep tember 1929 by arktisk næringsdrift, and named after anders kristian orvin [1889–1980], a geologist who worked for nsiu in spitsbergen and east greenland, and was director of norsk polar institut from 1958 to 1961. orvin was the first to land at this point. (orvinlia, orvin-lia.) oscar wisting bjerg 73ø-578 (73°46.2´n 27°47.0´w; map 4). mountain 2512 m high on the ne side of gerard de geer glet scher, named by høygaard and mehren in 1931 as oscar wistings fjell. the name appears to have been applied originally to a mountainous region 20 km ne of the present position. oscar adolf wisting [1871–1936] took part in the norwegian antarctic expedition to the south pole and the flight of the ‘norge’ with ellsworth and amundsen. osthytta – see østhytta. ostreaelv 70ø-105 (70°31.5´n 22°48.8´w; map 4). river in se jameson land west of kap stewart, named by alfred rosenkrantz during lauge koch’s 1926–27 expeditions in the form ostrea elv after the fossil oysters. it has also been called slate river. oststation 71ø (c. 71°03´n 24°15´w). locality in west jameson land, the site of alfred wegener’s 1930–31 eastern scientific sta tion, originally a wooden house. fuchs (1984) mistakingly identified lauge koch’s gurreholm station as this building. the german station was originally put up with the help of greenlanders from scorsbysund, who gave the locality the name tyskit nunaat; the station appears to have been dismantled after it closed in 1931. oswald heer hytten – see kap oswald heerhytten. oswald heer klinter 71ø-187 (71°28.0´n 24°18.6´w). low cliffs on the east side of schuchert dal. named during lauge koch’s 1936–38 expeditions by hans stauber after oswald heer. see also kap oswald heer. otocerasdal 73ø-50f (73°58.7´n 21°23.1´w). minor valley on the north slope of stensiö plateau, nw hold with hope, draining into blåelv. so named by eigil nielsen during the 1931–34 treårseks pedi tionen for the finds of fossil ‘otoceras’. otto johnsenvika 73ø (73°02.2´n 23°00.0´w). broad, open bay on the north coast of geographical society ø, se of robertson ø. used only on nsiu maps (lacmann 1937), and named after otto johnsen [b. 1901], a norwegian hunter who wintered in east greenland from 1929 to 1931 and 1932 to 1934. ottostrand 75ø (75°37.0´n 19°30.1´w). norwegian hunting station south of haystack on the east coast of hochstetter forland, one of john giæver’s main stations built in 1932. it was manned in the periods 1932–34, 1938–39 and 1948–53. the name commemorates the norwegian hunter otto johnsen. the station was also known as kolstad, and occasionally as norske roseneath to distinguish it from the danish hunting station mønstedhus, also called danske roseneath. overgangsdal 71ø-302 (71°39.2´n 24°40.5´w; map 5). valley on the north side of the front of bjørnbo gletscher, close to the boundary between crystalline and sedimentary rocks (overgang = transition). named by enrico kempter during lauge koch’s 1956– 58 expeditions. overkørslen 71ø (71°33.7´n 22°33.0´w). name sometimes used for the low col between inner nathorst fjord and carlsberg fjord, an easy sledge route. overkørslen 76ø-235 (76°46.2´n 18°37.9´w). low col east of dan markshavn, so named by the 1906–08 danmark-ekspedi tionen. this was the pass used by sledge parties proceeding northwards from danmark havn to lay out depots. (overkörslen.) oxford gletscher 71ø-369 (71°32.8´n 25°16.7´w; map 5). glacier in the south stauning alper, draining south into the east end of nordvestfjord. named by the 1962 oxford university expedition, which undertook survey work on the glacier. oxford university is one of the world’s oldest and most prestigious universities, whose origins go back to the early 12th century. uranus glacier has also been used. p p.k. larsen pynt 76ø-88 (76°40.8´n 18°30.6´w). south cape of renskæret, south of danmark havn. named by the 1906–08 danmark-ekspeditionen as p.k. larsens pynt. origin unknown. pad lochan 72ø (c. 72°14´n 23°55´w). name used by dundee university expeditions between 1968 and 1974 for a temporary water pool between mestersvig airfield and langdyssen. pain de sucre 70ø (70°43.4´n 25°58.9´w). isolated nunatak in charcot gletscher, east milne land. the name was used in the report by parat & drach (1934), and presumably derives from its colour and shape (pain de sucre = sugar loaf ). palasip qammavaajua [ferslew pynt] 70ø-305 (70°29.3´n 21°58.6´w). cape on the west side of ittoqqortoormiit [scores 273 bysund], close to ferslew pynt. recorded by the 1955 geodætisk institut name registration, the name means ‘the priest’s hunting place’, and was the locality where the settlement’s first priest, sejr abelsen, lay in wait while seal hunting. (palasip qámavâjua.) palasip qámavâjua – see palasip qammavaajua. palatinus 72ø (c. 72°05´n 25°05´w). mountain 2600 m high in the north stauning alper, nw of korsspids at the head of cavendish gletscher. climbed on 26 july by sandro pucci’s 1984 expedition. paletten 76ø-339 (76°13.8´n 26°21.6´w; map 4). group of nuna taks in sw dronning louise land. the name was given by the 1952–54 british north greenland expedition because the different coloured rocks forming the nunataks seemed to bare some resemblance to an artist’s palette. palisaderne 72ø-325 (72°32.4´n 24°11.0´w). peninsula on the north side of holm bugt, sw traill ø. the name was proposed by søkortarkivet in 1956–57 following surveying of the channel through vega sund as an alternative approach for ships en route to nyhavn (palisaderne = the palisades). palisaderne 80ø-40 (80°33.6´n 21°29.3´w; map 4). range of mountains on the west side of sødalen. named by eigil nielsen during the 1938–39 mørkefjord expedition for its appearance. palldal 71ø (71°32.8´n 24°11.0´w). valley in nw jameson land draining in schuchert dal, the present major paars dal. so named during l. koch’s 1936–38 two-year expedition by hans stauber (stauber 1940), because it was the winter pony route to ørsteddal used by pall pallson of reykjavik, who looked after the expedition ponies. palnatoke bjerg 74ø-358 (74°34.3´n 20°32.2´w; map 4). mountain 1056 m high in nw wollaston forland. the name was proposed by the place name committee in 1939 to replace a suggestion by wolf maync and andreas vischer, and commemorates the danish saga hero, palnatoke (toke palnessøn), founder of jomsborg, and the most celebrated of the joms vikings. (palnatokes bjerg.) palnatokeelv 74ø (74°31.6´n 20°34.9´w). river draining the slopes of palnatoke bjerg, north of zackenberg forskningsstation. the name is used as a reference locality by visiting scientists. pannekaka 73ø (73°07.3´n 22°43.7´w). small skerry off the coast of east ymer ø, so named on the 1932a nsiu map for its pancake-like shape. panorama lake 76ø (76°14.9´n 18°45.6´w). lake on store kolde wey where sampling was undertaken for phytoplankton studies (cremer et al. 2005). panoramafjeld 73ø-396 (73°31.4´n 25°19.0´w; map 4). mountain in east andrée land, west of the mouth of grejsdalen. it was named by erdhardt fränkl during lauge koch’s 1948–50 expeditions for the view. panoramaø 79ø-15 (79°18.8´n 19°08.7´w; map 4). small island near ne lambert land. named by the 1938–39 mørkefjord expedition, possibly by svend sølver who took a series of photographs here. panoramic peak 72ø (72°06.5´n 24°34.5´w; map 5). peak 1771 m high on the east side of bersærkerbræ, north stauning alper, climbed by the 1967 toni gobbi expedition. pap of cumbrae 71ø (71°57.2´´n 25°11.7´w; map 5). peak 1885 m high in the stauning alper, in the upper reaches of sefström glet scher. climbed by the 2001 scottish mountaineering club expedition, and so named for its nipple shape. parachute ponds 71ø (71°20.6´n 24°48.8´w). series of small ponds on the west side of the river draining holger danske briller. so named during the 1962 oxford university expedition (hall 1963, 1966) because six parachute loads of food and equipment were dropped here by dc3 on 19 july 1962. paradigma pass 71ø (71°42.3´n 22°37.9´w). name used by trümpy (1969) for the col north of paradigmabjerg, wegener halvø, and used in his desciption of ammonites collected during lauge koch’s 1958 expedition. paradigmabjerg 71ø-88 (71°41.8´n 22°37.3´w). mountain on wege ner halvø. so named by arne noe-nygaard during the 1931– 34 treårsekspeditionen as mt. paradigma because it produced a continuous geological sequence. paradisdal 73ø-634 (73°05.4´n 27°15.0´w). valley on the east side of kjerulf fjord near its junction with inner kejser franz joseph fjord. so named during the 1931–34 treårsekspeditionen by ove simonsen because of the rich vegetation and pleasant grassy slopes, which are in great contrast to the vertical and barren walls of nearby kejser franz joseph fjord. inuit ruins near the coast were excavated by j.m. wordie’s 1929 expedition. paradisdalen 73ø (73°05.8´n 27°18.2´w). name often used for the norwegian hunting hut at paradisdal, nw suess land. it was originally known as rendalshytta. paradisklippe 76ø-322 (76°41.8´n 24°15.9´w; fig. 21). long cliff on the north side of borgjökel, below himmerland hede, dron ning louise land. named by the 1952–54 british north greenland expedition. paradisterrasse 70ø-437 (70°30.0´n 29°24.3´w). plateau area between døde bræ and rolige bræ with an imposing view. named by laurent jemelin during the 1967–72 ggu scoresby sund expeditions for the biblical association with nearby arken and arrarat. paralleldal 73ø-79 (73°30.0´n 23°12.6´w; map 4). valley crossing central gauss halvø, trending approximately e–w and nearly following latitude 72°30´n. named by lauge koch’s 1929–30 expeditions in the form parallel valley (parallel = line of latitude). it has sometimes been used in the forms western parallel valley and eastern parallel valley for the parts draining respectively west and east (säve-söderbergh 1934). norwegian hunters have used the name tromsdal. (paralleltal.) paralleldalhytten 73ø (73°30.4´n 23°40.1´w). norwegian hunting hut built in october 1930 by arktisk næringsdrift on the south coast of gauss halvø, north of the mouth of paralleldal. it was originally known as dalheim. parat kløft 70ø-32 (70°44.6´n 25°31.3´w). ravine south of charcot havn, east milne land, named during the 1931–34 tre års ekspeditionen by hermann aldinger in the form paratschlucht, for maurice parat of the university of paris. he was a member of j.b. charcot’s 1933 expedition which visited the region, and was one of those drowned in the wreck of the pourquoi pas? in 1936. pariserøerne 78ø-8 (78°24.9´n 19°00.0´w; maps 1, 4). island group on the east side of jøkelbugten. named by the duke of orléans in 1905 as ile de paris, after the capital city of france, and possibly also after his father philippe d'orléans who was compte de paris. the 1906–08 danmark-ekspeditionen transferred the name to a group of islands 20 km west of the position estimated by the duke of orléans. (pariseröerne, pariser islands.) parker øer [islantit] 70ø-227 (70°43.4´n 21°29.8´w). two small islands off the coast of south liverpool land. named by william scoresby jr. in 1822 as parker island after a friend, charles parker. parkers piece 71ø-367 (71°56.6´n 25°22.7´w; map 5). ice plateau between pembroke kuppel and snetoppen, stauning alper. named by the 1963 cambridge university expedition after par ker’s piece, an open space south of emmanuel college in cam bridge commemorating edward parker, to whom it was leased in 1587. on published geodetic institute maps the name has been misplaced to the west of the glacier sometimes known as scorpio glacier. recent appoved name lists omit the ‘s’ such that the name becomes incorrectly ‘parker piece’. parkinson bjerg 73ø-60 (73°45.5´n 22°38.0´w; map 4). mountain in east hudson land, named by lauge koch’s 1929–30 expeditions as mt. parkinson after one of the geologists of wordie’s 1929 cambridge expedition, mark mervyn leofric parkinson. (parkin son fjellet, parkinsonberg, parkinson-berge.) parnas 71ø-347 (71°25.0´n 23°18.8´w). mountain in north 274 jameson land 1249 m high, adjacent to and higher than olympen. named during lauge koch’s 1958 expedition by john h. callomon after mount parnassus or óros parnassos in central greece. within sight of delphi and sacred to the dorians, it was favoured by roman poets as the home of the muses. pasdalen 74ø (74°26.7´n 19°51.5´w). name used by andreas vischer (vischer 1943) in a report on 1937 field work, for the valley issuing from a pass in the centre of eastern wollaston for land. pasdalshuset 71ø (c. 71°46´n 22°57´w). norwegian hunting hut built by helge ingstad’s expedition in 1932–33 at the mouth of solfaldsdal, fleming fjord. disappeared. it has also been known as syveren, mellemhuset and funkis. paselv 71ø-73 (71°11.1´n 22°30.7´w). river in northern klitdal draining into carlsberg fjord. so named during lauge koch’s 1926–27 expeditions by alfred rosenkrantz and tom harris as r. pass elv, because its course runs close to the pass in the valley between carlsberg fjord and hurry inlet. pasfjord 71ø-116 (71°16.0´n 21°56.5´w). short fjord west of kap vidar, east liverpool land. so named during the 1931–34 treårs eks peditionen by laurits bruhn because a pass in the valley at the head of the fjord leads over to carlsberg fjord. pashuset 74ø (74°25.0´n 20°19.1´w). danish hunting hut at the mouth of permdal, wollaston forland, a valley which leads to kuppelpasset. it was built by nanok in august 1933. pashytten 76ø (76°35.8´n 18°44.7´w). danish hunting hut on the east side of store koldewey, where a low pass leads over to berg fjord, built by nanok in august 1933. it is officially known as berg fjordhytten, and has also been known as yderhytten. pashytten 77ø-80 (c. 77°01´n 20°01´w). danish hunting hut ne of sælsøen, built by nanok in the spring of 1938. named for its position on the route to passet in central slædelandet. it has also been known under the names trekronerhytten, schultzhytten, hval sletten and slettehytten. passagegletscher 72ø-290 (72°49.0´n 28°16.4´w; map 4). glacier in south goodenough land on the west side of agassiz dal, used by john haller as a route westwards into the nunataks during lauge koch’s 1953 expedition. passagehøje 73ø-55 (73°53.9´n 22°11.2´w). mountain range rising to about 900 m west of loch fyne, named during lauge koch’s 1929–30 expeditions by helmar g. backlund as passage hills or passage berge. they had originally been called devon hills. (pas sage höhen.) passagen 71ø-181 (71°25.7´n 22°55.1´w; map 4). valley west of carlsberg fjord providing an easy connection to pingel dal. the name was one of a group of names given by the place name committee in 1939 to replace proposals by hans stauber. passe på 71ø (71°08.6´n 26°28.8´w). summit 2013 m high on the corner between edward bailey gletscher and catalinadal, ren land. climbed and named by the 2007 west lancashire mountain eering group expedition. passet 77ø-110 (77°07.5n 19°47.3´w). col on the sledge route through slædelandet, ne of mørkefjord station. named by the 1938–39 mørkefjord expedition. pasterze 74ø-67 (74°41.0´n 22°36.3´w; map 4). glacier west of the head of tyrolerfjord. so named during karl koldewey’s 1869–70 expedition by julius payer because of its azure blue colour and purity (payer 1876), and after the glacier of the same name in austria. see also grossglockner. pasterze (or pasterzenkees) is the largest glacier in the eastern alps. the original description of the east greenland glacier is that it seemed to be formed from five large tributaries, including the present copeland gletscher and kløft gletscher. the ice is considered by flint (1948) to have significantly retreated between 1869 and 1937. on the 1932 geodætisk institute 1:1 million scale map the main glacier was named gerda gletscher, and pasterze was applied to the present copeland glet scher. the valley was first explored by louise a. boyd in 1937. (pasterze glacier.) paul stern land 70ø-388 (70°24.0´n 29°29.0´w; maps 3, 4). large nunatak area west of vestfjord gletscher. named by eduard wenk after paul stern, a swiss geologist who worked in east greenland during the 1955–1958 lauge koch expeditions, and was the first to reach this nunatak. he died on 20 july 1959 in an accident on the winterstock in the urner alps. (poul stern land.) payer dal 74ø-340 (74°45.5n 20°17.4´w). major valley on south kuhn ø. named by wolf maync and andreas vischer during lauge koch’s 1936–38 expeditions (maync 1947). see also payer tinde. payer gletscher 73ø-637 (73°07.6´n 26°27.4´w). glacier in north suess land, west of payer tinde. named during the 1931–34 treårsekspeditionen by ove simonsen. the glacier used by payer during his ascent in 1870 was solklargletscher. see also payer tinde. (payers gletscher.) payer land 74ø-145 (74°30.0´n 22°30.0´w; maps 2, 4; fig. 15). area between wordie gletscher and tyrolerfjord, largely covered by an ice cap reaching an altitude of 1700 m. named by lauge koch’s 1929–30 expeditions to commemorate the observations by julius payer in this region in 1869. see also payer tinde. payer tinde 73ø-506 (73°07.6´n 26°21.6´w; map 4; figs 35, 68). mountain 2320 m high in north suess land, named by karl koldewey’s 1869–70 expedition as payerspitze or payer spitze for julius johannes ludovicus payer [1842–1915], an austrian army officer and polar explorer who was co-leader of the expedition. julius payer and ralph copeland climbed in august 1870, via solklargletscher, to the ice plateau ne of payer tinde; from here they had the first view of inner kejser franz joseph fjord and petermann bjerg. it has been claimed that the ascent of payer tinde in 1870 inaugerated arctic mountaineering (odell 1943), but john haller and wolfgang diehl who climbed payer tinde in 1952 found no evidence of a previous ascent (j. haller, personal communication). (payer peak, payers fjeld.) pebermyntefjeld 70ø-442 (70°29.2´n 28°57.5´w). mountain 1680 m high between rolige bræ and vestfjord. so named by laurent jemelin during the 1967–72 ggu scoresby sund expeditions because it is built up of red and white layers reminiscent of a peppermint. peder andresenvika 73ø (73°00.8´n 22°40.6´w). open bay on the north coast of geographical society ø, south of tveholmen. used on the nsiu maps of lacmann (1937), and named after peder andresen [b. 1891], a norwegian who was captain of the sæl barden on its 1934 voyage to east greenland. pederpynt 72ø-272 (72°52.8´n 24°49.2´w). minor cape on ne ella ø. named by john w. cowie during work carried out from 1949 to 54 on lauge koch’s geological expeditions, after peter j. adams, the british geologist who worked with him. pegasus gletscher 71ø-328 (71°44.4´n 25°15.3´w). glacier in the south stauning alper, a minor branch of bjørnbo gletscher. named by john hunt’s 1960 expedition as pegasus glacier after the constellation. pelion 71ø-405 (71°27.8´n 23°19.8´w). mountain 1200 m high in northern jameson land. named by katherina perch-nielsen during the 1967–72 ggu scoresby sund expeditions after pélion or óros pílion, a mountain chain in thessaly, greece. in greek mythology it was the home of centaurs. the name is in keeping with the features named olympen and parnas in the same region. pembroke kuppel 71ø-364 (71°56.1´n 25°21.3´w; map 5). snow dome about 2710 m high on the east side of spærregletscher, stauning alper. climbed by the 1963 cambridge university expedition on 8 august, and named after pembroke college, cam bridge, founded in 1347 by the countess of pembroke. pemmikanelv 76ø-300 (76°56.4´n 20°05.2´w). river in south ger mania land draining into slambugten east of hvalrosodden. so 275 named by the 1938–39 mørkefjord expedition because depots of pemmican were made here. pemmikankløft 76ø-303 (76°57.4´n 20°04.4´w). depression between østre skanse and vestre skanse occupied by pemmi kanelv, south germania land. named by the 1938–39 mørkefjord expedition. pendelbua 74ø (74°36.7´n 18°23.9´w). hunting hut on the south side of lille pendulum, built in the summer of 1921 for østgrøn landske fangstkompagni when it was known as kap desbrowe hus. it was repaired by the hird expedition in 1928, who subsequently described it as a norwegian hut under the names kap jona or pendelbua. (pendulumhytta.) pendulum øer 74ø-14 (74°39.0´n 18°41.0´w; maps 2, 4). island group off ne wollaston forland, made up of sabine ø, lille pen dulum, bass rock and hvalros ø. so named during douglas clavering’s 1823 expedition as the pendulum islands, because edward sabine swung the pendulum on the largest of the islands (sabine ø). sabine (1825) attributed the collective name to the officers and seamen of the griper. (pendulum-inseln, the pendu lum islands, pendulumön, pendulum øerne.) pendulumstrædet 74ø-3 (74°39.5´n 18°38.5´w; maps 2, 4). strait between sabine ø and lille pendulum, named by karl koldewey’s 1869–70 expedition as pendulum strasse. this may correspond to scoresby’s 1822 placing of gael hamke bugt. (pendulumstrasse, pendulumsundet, pendulum straits.) peninsola italica – see savoia halvø. penthievre fjord 77ø-9 (77°35.0´n 19°45.8´w; maps 2, 4). ne branch of skærfjorden, south of stormlandet. named by the duke of orléans in 1905 as fiord penthièvre after a branch of his family. his great-great-grandmother was louise-marie adélaide de bour bon-penthièvre. (penthièvres fjord.) perisphinctes ravine 74ø (74°45.4´n 19°58.2´w). ravine in se kuhn ø, named by maync (1947) for the finds of fossils during the 1936–38 two-year expedition. perka hytta 75ø (75°55.2´n 20°21.8´w). norwegian hunting hut built by john giæver’s expedition in august 1932 in the small bay known as pollen, south of the mouth of bessel fjord. it is also known as pollenhytta. perlehuset 70ø (70°47.0´n 24°08.5´w). name used for an inuit house ruin on the coast of jameson land 7 km south of falsterelv excavated in 1982 and 1983 (sandell & sandell 1985). it has yielded a very large collection of ornaments carved from bone and slate representing seals, birds and bears. danish archaeologists use the term ‘perle’ (= pearl) for ornamental objects made of different materials. permdal 71ø-180 (71°34.0´n 22°40.5´w; map 4). valley at the head of nathorst fjord. so named during lauge koch’s 1936–38 expeditions by hans stauber, presumably because the valley is formed in permian rocks. permdal 74ø-151 (74°23.8´n 20°10.1´w). valley in west wollaston forland, so named by h. frebold during the 1931–34 treårseks peditionen after the permian rocks. it has also been called zech stein dal. permklippen 72ø-224a (72°09.5´n 23°45.7´w; maps 4, 5). cliff about 100 m above sea level on the west side of the mouth of mesters vig. it was originally named bütlers klippe. named by prospecting teams associated with lauge koch’s 1948–49 expeditions. the rocks are of permian age. permpasset 74ø-350 (74°23.5´n 19°46.6´w). pass at the north end of blæsedalen, central wollaston forland. so named during lauge koch’s 1936–38 expeditions by wolf maync and andreas vischer for the outcrops of permian dolomites (maync 1947). pernaryggen 70ø-52 (70°43.2´n 25°24.3´w). ridge east of kronen and nw of kap leslie, east milne land, named during the 1931– 34 treårsekspeditionen by h. aldinger as pernarücken or perna rücken, for the fossil ‘perna’. perspective ridge 70ø (70°31´–70°47´n 22°15´w). as seen from hurry inlet the west side of liverpool land rises to what appeared to william scoresby in 1822 to be a level ridge 1500–2000 feet high, which he named for its form and appearance. the feature is marked, but not named, on scoresby’s chart and the name has not survived. it approximately corresponds to the present nukkaatsoq, heksefjeldet, gaffelfjeld and søbjergene. peschel island 75ø 76ø (76°06.0´n 21°08.0´w). name used in the english edition of koldewey’s 1869–70 narrative (koldewey 1874), for the landmass between bessel fjord and roon bugt which he payer tinde andrée land ättestupan kejser franz joseph fjord frænkel land fig. 68. looking west along kejser franz joseph fjord, with the 1300 m high cliff ättestupan on frænkel land to the left, payer tinde (2320 m high) on the skyline to the right, and andrée land in the background. 276 believed to be insular, and of which kap peschel is the ne cape. it corresponds to the present ad. s. jensen land. see also kap peschel. peter elv 72ø-216 (72°06.5´n 24°02.7´w; map 5). river on the north side of nedre funddal, north scoresby land, which joins with ping elv to form storm p. elv. named by prospecting teams associated with lauge koch’s 1948–49 expeditions. ‘peter & ping’ was the name of a cartoon series created by the danish artist storm p(etersen). (peters elv.) petermann bjerg 73ø-505 (73°05.4´n 28°37.1´w; maps 3, 4; figs 65, 69). mountain 2970 m high in west frænkel land, named by koldewey 1869–70 as petermanns spitze. it was first seen from the ice cap ne of payer tinde in august 1870, and described as an ice pyramid about 3300 m high, which could only be honoured by the name ‘petermann’. august heinrich petermann [1822–78] was a german geographer, a promoter of polar exploration, and publisher-editor of petermanns geographische mitteilungen. petermann was the driving force behind both the first and second german polar expeditions. a.g. nathorst in 1899 mistook a lower peak (now nathorst tinde) for petermann bjerg. the first ascent was made on 15 august 1929 by the cambridge expedition led by j.m. wordie, the second ascent by j. haller and w. diehl on 9 august 1951. (petermann peak, petermann fjeld, petermann point, peter manns bjerg, petermanns topp.) peters bugt 75ø-11 (75°18.0´n 20°08.0´w; map 4). bay on the sw side of hochstetter forland. named by karl koldewey’s 1869–70 expedition as peters bai, after wilhelm karl hartvig peters [1815– 1883], who wrote one of the zoological sections for koldewey’s expedition narrative. peters was a physician and zoologist who travelled in southern africa and madagascar (j. løve, personal communication 2010). (peters bay, petersbugt, perbugten.) peters bugt sø 75ø-112 (75°18.6´n 20°01.8´w). small lake on the east side of peters bugt. the name was first used by the 1976 swedish-danish east greenland expedition that had core-sampled the lake bottom sediments (see also björck et al. 1994). petersbugthytten 75ø-100 (75°20.1´n 20°11.8´w). danish hunting hut on the north side of peters bugt, hochstetter forland, built by nanok in august 1930. it has also been called bundhytten and nummer 1 hytten. (peters bugt hytten.) petersryggen 71ø-252 (71°57.6´n 23°51.8´w; map 5). mountain ridge in the werner bjerge on the east side of østre gletscher. named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk. it was climbed by bearth in 1953, and may have been named after him. (peters bjerg.) petrahytten 73ø (73°38.9´n 23°10.5´w). norwegian hunting hut on the north side of moskusoksefjord, 12 km se of hoelsbu, built in august 1932 for arktisk næringsdrift. it was originally known as først hytten, and acquired its present name when levin and petra winther took over the hoelsbu terrain in 1939. petra winther spent three years at hoelsbu with her husband levin from 1939 to 1942 (winther 1980). the hut has also been known as røiskattlia. petrol lochan 72ø (c. 72°14´n 23°55´w). name used by the uni versity of dundee expeditions between 1968 and 1974 for a temporary pool between mestersvig airfield and langdyssen. petters deep 77ø (c. 77°09´n 23°38´w). cove on the former north coast of britannia sø, dronning louise land, now concealed by the advance of britannia gletscher; a diesel generator supplied by a firm named ‘petter’ was lost here when a pontoon capsized near the 1952–54 british north greenland expedition base camp (simpson 1957). pevensey fjeld 71ø-338 (71°42.6´n 24°55.6´w; map 5). mountain 1811 m high on the north side of bjørnbo gletscher, south stau ning alper. first climbed by john hunt’s 1960 expedition, and named pevensey after pevensey castle, sussex, a norman castle dat ing from c. 1080. peveril 72ø (72°07.0´n 24°34.3´w; map 5). traversed by the 1982 sheffield university expedition, this mountain peak is described as the ‘bunny’s ears’ between arundel gate and beaumaris fjeld on the east side of bersærkerbræ, stauning alper. pfahl – see pælen. phecolitplateau 71ø-131 (71°05.7´n 22°04.5´w). plateau on the north side of inner storefjord, central liverpool land. so named by helge g. backlund during the 1931–34 treårsekspeditionen. probably derived from the geological term phacolith, a minor intrusion in folded rocks. philipshorn 72ø-116 (72°22.6´n 25°55.8´w; map 4). mountain on the south side of forsblad fjord, so named during the 1931–34 treårsekspeditionen by eugéne wegmann. it is recorded as not to have been named after a specific person, but was to be considered to honour the noted glaciologist professor philipp, or the philip bræen on spitsbergen. (philiphorn, philippshorn.) phillips' point 70ø (70°34.4´n 22°34.7´w). named by william scoresby jr. in 1822 after one of his two partners in the baffin. the name does not appear on his map, but the latitude and longitude in the appendix show it to be a point on the west side of hurry inlet, possibly that now known as albuen. phynoldsbjerg 73ø-686 (73°30.0´n 26°24.4´w). mountain on the sw side of djævlekløft. named during lauge koch’s 1949–51 expeditions by john haller, apparently after the son of the tele graphist (aage de lemos) on ella ø. pianofirn 72ø-248 (72°17.1´n 24°36.6´w). glacier in the north stauning alper between skjoldungebræ and syltoppene, named by erdhardt fränkl during lauge koch’s 1950–51 expeditions. the glacier has a step-like profile with black and white stripes said to be reminiscent of the keys of a piano. pic andersen 71ø (71°56.2´n 25°45.5´w). mountain about 2450 m high on the east side of prinsessegletscher. named and first climb ed by claude rey’s 1968 expedition. pic andrée georges 71ø (71°57.3´n 25°47.7´w). mountain on the east side of prinsessegletscher. first climbed by claude rey’s 1968 expedition; the climbing party included jean-louis georges. pic brian roberts 70ø (70°46.7´n 25°59.3´w). mountain 1691 m high on east milne land on the north side of charcot gletscher. the name was used by parat & drach (1934), and was named after brian birley roberts [1912–78], leader of the 1933 cambridge expedition that had been transported to and from greenland on the pourquoi pas? roberts also took part in the 1934–37 graham land expedition (antarctica), and after joining the scott polar research institute in 1946 was closely involved in antarctic affairs, notably the 1959 antarctic treaty. pic de gerlache 78ø-2 (78°36.3´n 21°27.7´w; map 4). pronounced peak in nørre biland, the northern part of hertugen af orléans land. it was named by the duke of orléans in 1905 for the belgian polar explorer, adrien victor joseph baron de gerlache de gomery [1866–1934]. adrien de gerlache led the belgian antarctic expedition 1897–99, the first to over-winter in the antarctic, and com manded the belgica in 1905 for the duke of orléans. he also took part in further arctic expeditions, in 1907 to nova zemlya and 1909 to east greenland, spitsbergen and frans josef land. the 1906–08 danmark-ekspeditionen was unable to identify the original peak, but considered the name should be preserved and placed it on a prominent mountain 912 m high on the north side of gam mel hellerup gletscher. on some maps the name is misplaced westwards to a slightly higher but less prominent peak. pic du pourquoi-pas? 70ø (70°41.0´n 26°02.1´w). mountain 1643m high on the south side of charcot gletscher. named by parat & drach (1934). see also pourquoi pas tinde. pic flotard 71ø (71°52.0´n 25°48.5´w). peak 2200 m high on the west side of prinsessegletscher. named and first climbed by claude rey’s 1968 expedition. pic ludovica 71ø (71°55.3´n 25°45.0´w). mountain about 2400 m high on the east side of prinsessegletscher. named and first climb 277 ed by claude rey’s 1968 expedition. piccadilly 72ø (72°08.5´n 24°31.7´w; map 5). mountain 1692 m high on the east side of bersærkerbræ, north stauning alper. first climbed by the 1963 imperial college expedition, and named after the london street, one of the two ancient highways leading west out of london. pictet bjerge 72ø-1 (72°04.5´n 23°23.0´w; map 4). mountain range on the south side of davy sund. named by william scoresby jr. in 1822 as cape pictet, for marc auguste pictet [1752–1825], who held the chair of natural philosophy at geneva from 1786 to 1825. scoresby’s cape was evidently a mountain, and nathorst (1901) transferred the name to the mountains behind the present kap syenit. (pictet bjerg, pictet mountains, mt. pictet, pictetfjella.) pictetbjerghytten 72ø (72°07.5´n 23°28.6´w). name commonly used for the norwegian hunting hut built by the møre expedition in august 1930 at the foot of pictet bjerg, on the south side of davy sund. it was originally known as jostein, and has also been known as segldalen and bjørnebu. piggøyra 72ø (72°40.5´n 22°01.9´w). peninsula on se geographical society ø, a little west of kap mcclintock. so named on the nsiu maps of lacmann (1937) after its shape (pigg = spike). (piggöyra.) pilgrimsdal 72ø-475 (72°06.5´n 26°22.9´w). valley north of the nw end of furesø, nathorst land, named by hans zweifel during the 1954–55 lauge koch expeditions. this is a steep glacier-filled valley, not easy for a ‘sinner’ to climb (fritz schwarzenbach, per sonal communication 1996). pillen 76ø-289 (76°50.5´n 20°21.9´w). small island in north dove bugt between vindseløen and fugleø. so named by the 1938–39 mørkefjord expedition because of its pillar-like appearance. pimlico 72ø (72°09.5´n 24°42.2´w; map 5). mountain 1850 m high on the north side of dunottar gletscher, north stauning alper. first climbed by the 1963 imperial college expedition and named after the sw london district. pindsvinet 70ø-270 (70°01.8´n 23°21.9´w). mountain about 1730 m high on volquaart boon kyst, so named during the 1931–34 treårsekspeditionen by laurits bruhn for its spiked basalt pinnacles (pindsvin = hedgehog). ping elv 72ø-217 (72°06.2´n 24°03.5´w; map 5). river on the south side of nedre funddal, north scoresby land, which joins with peter elv to form storm p. elv. named by prospecting teams associated with lauge koch’s 1948–49 expeditions. ‘peter & ping’ was the name of a cartoon series created by the danish artist storm p(etersen). (pings elv.) pingel dal 71ø-49 (71°32.1´n 23°01.3´w; map 4). eastern of two large valleys which drain to the head of fleming fjord. named by g.c. amdrup’s 1898–1900 expedition as pingels dal, probably after the danish geologist chr. pingel and his son j.v. pingel. nor wegian hunters used the name fleming dal for the valley, and the hunting hut built in the valley in 1932 is known as both pingel dal hytten and fleming dal hytten. (pingels dale, pingeldal, pingel valley.) pingo dal 71ø-172 (71°47.4´n 23°49.1´w; maps 4, 5). valley south of the werner bjerge where a number of pingos are found. pingos are characteristic volcano-like mounds (fig. 70), here up to 30 m high. fritz müller who had studied them during lauge koch’s geo logical expeditions, argued strongly for the name, which replaced the name kristiern nielsen dal, rarely used but officially approved from 1937 to 59. a hut at the head of the valley close to lomsø, usually known as lommensø hytten, has sometimes been referred to as pingo dal hytten. pingoelv 71ø (71°47.4´n 23°49.1´w). informal name used by fritz müller during lauge koch’s 1954–55 expeditions for the river in pingo dal, north jameson land (müller 1959). pingo pass 71ø (71°47.7´n 24°13.1´w; map 5). name occasionally used for the pass at the head of pingo dal leading over to schuchert dal (e.g. schwarzenbach 1996). pinkfoot lake 77ø (77°40.1´n 20°42.2´w). lake in nordmarken, north of klægbugt, where pink-foot geese congregate. named by the 1987 irish expedition to northern east greenland. pinkfoot pond 71ø (71°46.4´n 23°24.9´w). name used in an orni thology report of the 1963 british east greenland expedition (hall & waddingham 1966) for a lake on the north side of ørsted dal. nests of the pink-footed goose were found here. pinnacle 71ø (71°55.2´n 24°58.0´w; map 5). mountain on the ridge lille petermann petermann bjerg løvebastion nordenskiöld gletscher fig. 69. view from goodenough land westwards across nordenskiöld gletscher to lille petermann (2709 m) and petermann bjerg (2970 m). løvebastion to the left has an appearance of a lion when viewed from nordenskiöld gletscher. 278 between storgletscher and dalmore glacier, central stauning alper. named by the university of dundee expedition which made the first ascent on 15 august 1968. (pinnacle peak, the pinnacle.) pinnacle mount – see murtinderne. pinnadal 70ø-51 (70°42.6´n 25°17.8´w). small valley on the east coast of milne land between kap leslie and charcot havn. named during the 1931–34 treårsekspeditionen by hermann aldinger as pinnatal or pinna-tal, after the fossil lamellibranch ‘pinna’. (pinna valley.) pisa 71ø (71°40.3´n 24°58.5´w; map 5). small rock peak 1350 m high on the north side of the junction of mercurius gletscher and bjørnbo gletscher. first climbed by james clarkson’s 1961 expedition, and so named because it resembled a large leaning tower. pissevache 73ø-426 (73°20.0´n 24°45.9´w). waterfall in north ymer ø, named by silvio eha during lauge koch’s 1947–49 expeditions after the waterfall of the same name in the rhone valley near martigny. piz coaz 71ø (71°53.9´n 25°27.2´w; map 5). narrow ridge reaching 1950 m high between spærregletscher and duart gletscher, stau ning alper. first climbed by the 1964 aac zürich expedition which named it for johanne coaz, a pioneer swiss climber who made the first ascent of piz bernina in 1850, the highest point in the engadine. the second climb of piz coaz was made by karl m. herligkoffer’s 1966 expedition, which called it rosenheimer spids. piz dominant 71ø (71°54.6´n 25°34.3´w; map 5). peak about 2370 m high on the west side of spærregletscher, stauning alper. first climbed and so named by the 1964 aac zürich expedition. the second ascent was by karl m. herligkoffer’s 1966 expedition. piz guarda monti 71ø (71°57.1´n 25°36.0´w; map 5). peak about 1840 m high on the west side of spærregletscher, stauning alper. climbed and so named by the 1964 aac zürich expedition. piz spescha 71ø (71°57.5´n 25°27.9´w; map 5). mountain 2210 m high east of spærregletscher. first climbed by the 1964 aac zürich expedition, which named it after father plàcidus spescha, a benedictine monk who made a number of pioneer climbs in the swiss alps between 1788 and 1824. it is identical with schöne aussicht. piz vadian 71ø (71°58.9´n 25°33.4´w; map 5). mountain 1640 m high east of spærregletscher, stauning alper. so named and first climbed by the 1964 aac zürich expedition. place concordia 70ø (70°43.2´n 25°56.9´w). broad, circular, flat area of charcot gletscher on east milne land. so named by parat & drach (1934), after the similarly named glacier confluences in the swiss alps. pladebjerg 73ø-702 (73°13.3´n 26°53.1´w). mountain in eastern frænkel land so named during lauge koch’s 1949–51 expeditions by john haller, because the entire mountain is made up of rocks rich in plates (= plade) of mica. pladen 74ø-283 (74°09.9´n 20°52.7´w). mountain on se clavering ø. the name originated from the wintering party at eskimonæs during the 1931–34 treårsekspeditionen (plade = slab, plate). herman andresenfjellet has also been used. pladen 76ø-56 (76°51´n 20°05´w; map 4). small elongate island in the north part of dove bugt. named by the 1906–08 danmarkekspeditionen (pladen = the slab). planck klippe 76ø-314 (76°57.9´n 24°15.9´w; map 4). cliff on the south side of admiralty gletscher, east of regnbue klippe, dron ning louise land. one of the names given by the 1952–54 british north greenland expedition for notable scientists, it commemorates the german physicist max [karl ludwig] planck [1858– 1947], noted especially for the quantum theory. plant hill 73ø-295 (73°55.0´n 22°11.7´w). mountain in east hudson land. it was named by gunnar säve-söderbergh during the 1931–34 treårsekspeditionen for the finds of fossil plants, which showed the rocks to be younger than had been first sup posed. pleinting bjerg 71ø (71°51.9´n 25°15.4´w; map 5). mountain on the south side of the head of roslin gletscher. climbed by karl m. herligkoffer’s expedition on 15 august 1966, and named after the small bavarian town of pleinting. (pleintingbjerg.) plinganser col 71ø (71°51.5´n 25°25.2´w; map 5). col between the upper part of duart gletscher and the upper basin of spærre gletscher. climbed and so named by karl m. herligkoffer’s 1966 expedition. plovjernet 70ø-77 (70°19.0´n 25°05.6´w; map 4). mountain on the east side of vikingebugt. so named by laurits bruhn during the 1931–34 treårsekspeditionen for its appearance (plovjern = plough share). pluto nunatak 72ø-295 (72°52.5´n 29°15.8´w). nunatak on the west side of nordenskiöld gletscher, where the danish air force catalina 853 ‘pluto’ dropped provisions on 23 july 1953 for the use of a geological exploration party. named by john haller, who reached the summit with other members of the party on 8 august 1953. podiet 70ø-451 (70°19.3´n 29°31.0´w). cliff in se paul stern land overlooking vestfjord gletscher. so named by w.e.a. phillips during the 1967–72 ggu scoresby sund expeditions because it rembles a podium. point ambler 70ø (70°50.3´n 26°04.6´w). summit on the north side of polkorridoren, milne land. climbed by the 2004 west lancashire scouts expedition. point hope 70ø (70°27.4´n 22°16.1´w). cape in south liverpool land a few kilometres east of kap hope, corresponding to the present basaltnæs. it is marked on maps in e. mikkelsen (1927), and appears in some accounts as falska kap hope. timber was left here in 1924 during the colonisation expedition, but later moved farther west to the present settlement. point jilly 72ø (72°06.2´n 24°54.9´w). prominent pinnacle on the north ridge of the mountain lambeth, stauning alper. it was climbed by the 1996 scottish mountaineering club expedition. point neurose 72ø (72°03.9´n 24°44.5´w). mountain at the head of schuchert gletscher, stauning alper. the position is somewhat uncertain, but is described in bennet (1972) as a short distance sw of royal peak. it was climbed by the 1961 bangor expedition pointe c. jacquemard 71ø (71°54.4´n 25°53.4´w). mountain on the west side of prinsessegletscher. named and first climbed by claude fig. 70. conical mound known as a pingo, observed in margrethe dal on gauss halvø in 1990. these ice-cored glacial features can be up to 500 m in diameter and 50 m high. in east greenland they are often developed on braided river plains (bennike 1998). 279 rey’s 1968 expedition. exact position uncertain. pointe humbert 71ø (71°52.9´n 25°52.0´w). mountain about 2100 m high on the west side of prinsessegletscher, north of gl. des violettes. named and first climbed by claude rey’s 1968 expedition. pointe michel gravost 72ø (72°11.8´n 25°11.9´w). peak on the north side of vikingebræ, north stauning alper, climbed by claude rey’s 1970 expedition. it was reported by bennet (1972) to be probably identical with one of the dreispitz. pointe d'argent 71ø (71°54.0´n 25°54.7´w). rock pillar about 2480 m high on the west side of prinsessegletscher, south of combe d’argent. named and first climbed by claude rey’s 1968 expedition. pointe de france 71ø (71°55.5´n 25°55.2´w). rock peak about 2210 m high on the west side of prinsessegletscher. named and first climbed by claude rey’s 1968 expedition. exact position uncertain. pointe des ours 71ø (71°54.3´n 25°53.2´w). mountain on the west side of prinsessegletscher. named and first climbed by claude rey’s 1968 expedition. exact position uncertain. polar bear lake 76ø (lake i, 76°14.2´n 18°43.8´w; lake iii, 76°14.0´n 18°46.2´w). three small lakes on store koldewey, of which lakes i and iii were sampled for phytoplankton studies (cremer et al. 2005). polarheimen 73ø (73°11.1´n 25°58.4´w). norwegian hunting hut on the ne coast of suess land, built by arktisk næringsdrift in july 1947. it has also been known as røiskattlia and gråkollen. polheim – see polhemsdalhytten. polhem dal 72ø-35 (72°36.2´n 25°16.1´w; map 4). major n–strending valley in east lyell land. named by a.g. nathorst’s 1899 expedition, probably after the steamer polhem used on the 1872– 73 expedition to spitsbergen. (polhem valley, polhems dal, polhem dalen.) polhemsdalhytten 72ø (72°26.7´n 25°28.9´w). norwegian hunting hut in forsblad fjord, 2 km west of polhem dal, built in september 1931 by the møre expedition. it was originally known as bærtun. (polheim, polhem dal hytten.) pollen 75ø-63 (75°55.6´n 20°21.7´w). bay sw of trums ø, south of the mouth of bessel fjord. the name originated from the wintering party at kulhus during the 1931–34 treårsekspeditionen, and is a norwegian word for a small bay. the hut in the bay, originally known as perka hytta, is sometimes referred to as pollenhytten. pollux 71ø (71°50.6´n 25°31.5´w; map 5). peak on the sw side of the upper basin of spærregletscher, very close to the summit castor. named by karl m. herligkoffer’s 1966 expedition, al though not climbed. see also pollux elv. pollux elv 70ø-182 (70°34.5´n 22°23.9´w). one of a pair of similar rivers in south liverpool land draining west into hurry inlet, the other being castor elv. named during the 1931–34 treårseks peditionen by laurits bruhn after the stars castor and pollux, which derive their names from the twins of greek mythology. pollux glacier 71ø (71°56.7´n 25°37.7´w; map 5). one of two minor tributaries to spærregletscher on its west side, named by james clarkson’s 1961 expedition. see also pollux elv. german mountaineering accounts use kleine sydney gletscher for the same glacier. polluxbjerg 71ø-297 (71°57.0´n 26°16.5´w; map 4). mountain about 2300 m high on the south side of the west end of furesø, opposite castorbjerg on the north side of the lake. the two mountains were named by hans zweifel during lauge koch’s 1954–55 expeditions after two mountains with similar names in wallis, switzerland. polypen gletscher 72ø-166 (72°27.9´n 22°10.4´w). glacier on eastern traill ø on the nw flank of ellemandsbjerge. named during lauge koch’s 1936–38 expeditions by hans p. schaub after the expedition motorboat polypen. see also polyphavn. polyphavn 71ø-111 (71°13.4´n 21°45.7´w). short fjord in east liverpool land south of trekanten, named during the 1931–34 treårsekspeditionen by helmar g. backlund as polyp havn after the motorboat polypen used during their explorations in 1933. the polypen was one of lauge koch’s largest motor boats. it was holed by german troops in the spring of 1943, repaired and used by the sirius sledge patrol, but damaged when picked up by the american patrol boat eastwind. it ended as a wreck at kap berghaus. polyptychiteselv 72ø-233 (72°28.1´n 22°45.1´w). river on east traill ø, draining south into mountnorris fjord. so named by des mond t. donovan during lauge koch’s 1949–50 expeditions for the important fossils found here. pony gletscher 76ø-328 (76°28.4´n 25°00.0´w; map 4; fig. 21). glacier in dronning louise land flowing se to vedel sø. named by the 1952–54 british north greenland expedition after the ponies which j.p. koch’s 1912–13 expedition used on their traverse of this glacier, and across the inland ice. poplar 72ø (72°10.2´n 24°40.8´w; map 5). mountain 1850 m high between dunottar gletscher and harlech gletscher, north stau ning alper. it was first climbed by the 1963 imperial college expedition, and named after the london borough. porfyrbjerg 71ø-95 (71°43.3´n 22°17.4´w). mountain in north canning land, named during the 1931–34 treårsekspeditionen by arne noe-nygaard as porfyrfjeld or porfyrfjeldet, after the por phy ritic volcanic rocks. porfyrdal 71ø-101 (71°41.1´n 22°16.3´w). valley in north canning land south of porfyrbjerg. named by arne noe-nygaard during the 1931–34 treårsekspeditionen as porphyry valley. porfyrryggen 72ø-397 (72°03.9´n 23°31.8´w). mountain ridge in north scoresby land between antarctic havn and jægerdal. named by hans kapp during lauge koch’s 1957–58 expeditions for the rock types. port arthur 76ø-148 (76°46.3´n 21°12.3´w; map 4). circular bay on the east side of daniel bruun land, so named by j.p. koch’s 1912–13 expedition for an apparent similarity with port arthur, a major port city in ne china. port arthurhytten 76ø-199 (76°45.9´n 21°05.3´w). danish hunting hut north of the mouth of port arthur on the se coast of daniel bruun land, built by nanok in august 1933. it is also known as spydodden. (port arthur hytten.) portalen 74ø (74°00.1´n 21°23.6´w). name used by eigil nielsen (1935) for a feature in the valley niviarsiaq (river 13) in north hold with hope, which resembles a gateway. porten 73ø-653 (73°36.4´n 24°41.2´w). valley in the high cliffs of south strindberg land. so named by th. johansen during the 1931–34 treårsekspeditionen because it has the appearance of a gateway (= port). porten 74ø-298 (74°37.5´n 20°52.4´w). mountain at the mouth of slettedalen east of the stream which provides a steep route up to the valley (porten = the gate). the name originated from the wintering party at kulhus during the 1931–34 treårseks peditionen. portfjeldet 80ø-44 (80°32.7´n 21°04.2´w). mountain on the south side of the mouth of sødalen. named by eigil nielsen during the 1938–39 mørkefjord expedition in the form portfjældet (port = gateway). portgletscher 73ø-611 (73°11.8´n 27°48.9´w). glacier on the south side of knækdalen, under which knækelven flows beneath an arch of ice, named by louise boyd’s 1933 expedition as arch glacier. the arched tunnel was 80 m long in 1933, and still existed in 1975. portmorænen 73ø-614 (73°13.0´n 27°57.8´w). moraine barrier across central knækdalen, formed by gregory gletscher when it was 9 km in advance of its present position. named by louise boyd’s 1933 expedition in the form gateway moraine, because 280 knækelven has cut a narrow opening in the moraine wall. posten 73ø-584 (73°57.6´n 24°18.8´w; map 4). mountain on the east side of waltershausen gletscher. so named by skaun & welde’s 1932 expedition, probably after ‘dagsposten’, the norwe gian newspaper which supported their expedition. posten 80ø-80 (80°03.5´n 20°12.0´w). mountain in south kron prins christian land, south of marmorvigen. so named during lauge koch’s 1952–53 expeditions by erdhardt fränkl, because of the superb view from the summit which would make it a ‘good site for a fortress or mountain hotel’. postkassen 70ø-129 (70°50.3´n 22°43.3´w). mountain west of the head of hurry inlet, named by alfred rosenkrantz during lauge koch’s 1926–27 expeditions in the form letter box mt. it is said to have been named for its shape. potamogetonsø 70ø-389 (70°58.6´n 27°43.1´w). small lake on c. hofmann halvø, south of rypenæs. the name was approved in 1961 at the suggestion of ulrik røen, and records finds of the water-plant potamogeton. pothorst bjerge 71ø-178 (71°35.0´n 23°39.6´w). mountain range south of the head of ørsted dal. the name was one of a group of names given by the place name committee in 1939 to replace proposals by hans stauber. it was given for the dane who the danish king christian i sent to greenland with didrik pining in 1476. poulsen nunatakker 76ø-151 (76°56.3´n 26°22.5´w; map 4). group of three isolated nunataks in west dronning louise land. named by the 1909–12 alabama expedition as poulsen’s nuna takker after georg poulsen, mate and member of the expedition who took part in the sledge journey to dronning louise land in april 1910. on recent official lists the name appears in the singular as poulsen nunatak. pourhelène 71ø (71°11.4´n 26°28.0´w). mountain 1909 m high in renland. climbed and named by the 2007 west lancashire moun ain eering group expedition. pourquoi pas tinde 70ø-90a (70°40.3´n 25°51.0´w; fig. 71). moun tain 1011 m high on se milne land. the earliest appearance of the name occurs in a report by parat & drach (1934) in the form pic de pourquoi pas?, and was originally applied to a 1643 m high mountain 7 km west (70°41´n 26°03´w) of that which now bears the name. the pourquoi pas? was a 3-mast ice-strengthened barque, built in st. malo in 1907 for jean-baptiste charcot and taken over by the state for the 1908–10 french antarctic expedition. it was subsequently used by charcot on numerous arctic voyages, including seven to the scoresby sund region, and was wrecked off iceland in september 1936. mt. rosenkrantz has also been used for this mountain. (pourquoi-pas tinden.) priener kalotte 71ø (71°53.3´n 25°24.0´w; map 5). mountain about 2100 m high on the sw side of duart gletscher, central stauning alper. climbed by karl m. herligkoffer’s 1966 expedition on 17 august, and named after the small town of prien on the chiemsee at the foot of the bavarian alps. priener spids 71ø (71°53.0´n 25°24.0´w; map 5). mountain on the west side of duart gletscher. climbed by karl m. herligkoffer’s 1966 expedition on 17 august. see also priener kalotte. (priener spids.) primula pond 71ø (71°41.3´n 23°43.9´w). name used in an orni thology report of the 1963 british east greenland expedition (hall & waddingham 1966) for a lake in ørsted dal, close to the mouth of pingo dal. it was named for the many flowers. primulabugt 73ø-602 (73°48.8´n 25°24.6´w). pronounced small bay on the west coast of strindberg land. the name first appears as a botanical reference locality in a report of the 1931–34 treårs ekspeditionen (gelting 1934). gunnar seidenfaden had collected specimens of ‘primula stricta’ here in 1929, at the time the north limit of primula in east greenland. primulaelv 70ø-134 (70°42.0´n 22°46.2´w). river on the west side of hurry inlet entering the fjord at constable pynt. named during lauge koch’s 1926–27 expeditions by alfred rosenkrantz, appearing first as primula river in the report by harris (1931), although without precise location. the name was given for the abundant flowering primula. (primula elv.) primulaklöft 70ø (70°42.0´n 22°46.2´w). name used by rosen krantz (1934) for the ravine in which primulaelv flows. prins axel nunatak 77ø-45 (77°15.8´n 24°17.6´w; map 4). large nunatak in north dronning louise land, named by the 1909–12 ala bama expedition as prins axel’s nunatak for prince axel chri stian georg of denmark [1888–1964], an uncle of frederik ix. he was a danish naval officer, who reached the rank of captain in 1918, and orlogskaptajn (lieutenant commander) in 1923 (j. løve personal communication 2009). prinsen af wales bjerge 69ø-36 (69°01.0´n 32°42.0´w). group of nunataks north of kangerlussuaq (68°35´n), named after the bri t ish crown prince [1894–1972], briefly edward viii in 1936, who was patron of the 1934 british trans-greenland expedition (lind say 1935). (prince of wales mtns.) prinsesse caroline-mathilde alper 80ø-50 (80°24.0´n 19°47.0´w; maps 1, 4). mountain range in east kronprins christian land, south of ingolf fjord. named by the 1938–39 mørkefjord expedition after the wife of the danish prince knud, patron of the expedition, for her always friendly interest. princess caroline-mathilde [1912–1995] was noted especially for her patronage of danish organisations in sydslesvig. (prinsesse caroline-mathildes alper.) prinsesse elisabeth alper 80ø-51 81ø-130 (80°48.0´n 18°45.0´w; maps 1, 4). mountain range north of ingolf fjord, trending ne– sw. named by eigil nielsen during the 1938–39 mørkefjord expedition after the danish princess elisabeth [b. 1935], daughter of fig. 71. this was the fourth ship of the same name, pourquoi pas?, all owned by the french polar explorer jean-baptist charcot. built in 1908, this 40 m, 455-ton barque was wrecked off iceland on 16 september 1936, with the loss of 39 crew and scientists; only one man survived. photo: kindly supplied by emilie thomassot, © centre de recherches pétrographiques et géochimiques, nancy, france. 281 prince knud. (prinsesse elisabeths alper.) prinsessegletscher 71ø-299 (71°57.0´n 25°50.5´w; map 5). major glacier on the south side of furesø. named by john haller following explorations during lauge koch’s geological expedition in 1954, probably for one of the three danish princesses, daughters of frederik ix. glacier du furesoe has also been used. prinsessen 77ø-128 (77°04.1´n 25°07.3´w; map 4; fig. 21). specta cular ice-covered mountain in nw dronning louise land. named by the 1951 british north greenland reconnaissance expedition after the patron of the expedition, then princess elizabeth, now queen elizabeth ii, queen of the united kingdom and the com monwealth. she succeeded to the throne on 6 february 1952. see also hertugen. prinsessen col 77ø (77°03.7´n 25°05.5´w). name used occasionally in expedition reports (simpson 1957) for the col immediately se of prinsessen, dronning louise land. proctor’s pinnacle 72ø (72°07.5´n 25°07.8´w; map 5). pinnacle 2350 m high at the corner of vertebrae and gully gletscher, stauning alper. climbed by the 1963 cambridge university expedition, which named it after the cambridge university officials known as proctors (pinnacolo di proctor.) productuselv 74ø (74°13.6´n 20°40.3´w). river on east clavering ø draining from the slopes of binucleus and trinucleus. named by wolf maync and andreas vischer during lauge koch’s 1936–38 expeditions, and used in the report by maync (1942) and on ams maps. fossil productids were found here. profilbjerg 71ø (71°37.2´n 22°56.2´w). name introduced by stauber (1942) for the sw peak of lille cirkusbjerg, south wegener halvø, where a profile was measured during lauge koch’s 1936–38 twoyear expedition. the name has been frequently used as a reference locality in geological literature (grasmück & trümpy 1969; hig gins 1986). profilbjerg 72ø-192 (72°07.6´n 24°06.9´w; map 5). mountain in north scoresby land, bounded to the south by nordre funddal and nedre funddal. named by prospecting teams associated with lauge koch’s 1948–49 expeditions. (profilbjerget.) profildal 71ø-421 (71°07.8´n 27°34.6´w; map 4). valley in se renland, east of rypefjord. so named by johan d. friderichsen during the 1967–72 ggu scoresby sund expeditions because of a well-exposed geological profile. profile ravine 73ø (73°30.7´n 23°15.8´w). ravine on the south side of sederholm bjerg, central gauss halvø. the name was used by gunnar säve-söderbergh during the 1931–34 treårsekspedi tionen, because geological profiles were measured here. profilfjeldet 80ø-42 (80°31.5´n 21°26.5´w). mountain on the north side of sødalen. named by eigil nielsen during the 1938–39 mørkefjord expedition as profilfjældet, because geological sections were measured here. promenadedal 74ø-135 (74°03.8´n 23°06.4´w; map 4). prominent valley on the south side of wordie gletscher, named by lauge koch’s 1929–30 expeditions in the form promenade valley. it is a long and wide valley and is easy walking terrain (promenade = promenade, parade). (promenade tal, promenadetal, gangdalen.) prometheus 71ø (71°44.7´n 25°25.3´w; map 5). mountain 2574 m high on the sw side of orion gletscher, south stauning alper. it was first climbed by james clarkson’s 1961 expedition, and named after the god of fire in greek mythology. proppen 70ø-401 (70°57.6´n 28°29.3´w; map 4). nunatak at the head of harefjord. so named by the 1963 geodætisk institute ex pedition because it blocked the flow of ice like a cork (= prop) in a bottle. prospekt dal 73ø-75 (73°36.0´n 22°38.7´w). small valley west of ankerbjerg on the north side of moskusoksefjord, named by lauge koch’s 1929–30 expeditions as prospect valley, because of a possible mineralisation prospect. norwegian scientists used vassdalen for the same feature. (prospekttal, prospectdal.) prospektfjeld 72ø-239 (72°19.6´n 22°40.8´w). mountain on se traill ø on the south side of mountnorris fjord. so named by desmond t. donovan during lauge koch’s 1949–50 expeditions. prædikestolen 70ø-412 (70°54.0´n 28°17.6´w). mountain 1271 m high south of inner harefjord. named during the 1967–72 ggu scoresby sund expeditions by heintz rutishauser for its resemblance to the mountain ‘kanzel’ in upper lauterbrunnen tal, swit zer land (prædikestolen = kanzel = pulpit). præstekravesø 70ø-434 (70°27.5´n 27°40.9´w). small lake on sw milne land. named during the 1967–72 ggu scoresby sund ex peditions by max fumasoli for the numerous ringed plovers (= præstekraver). prøvestenen 76ø-136 (76°25.6´n 26°45.7´w; map 4). nunatak in sw dronning louise land, so named by j.p. koch’s 1912–13 expedition after one of the coastal sea forts off copenhagen. (pröve stenen.) puchan glacier 72ø (72°04.4´n 25°02.9´w). name used by the 2007 smc east greenland expedition for the western upper branch of gullygletscher. puchwhitstinde 72ø (72°00.6´n 24°45.7´w; map 5). mountain 2339 m high on the east side of upper storgletscher, central stauning alper. climbed and named by the 2007 smc east greenland expedition. the name derives from a combination of personal names. puderne 70ø-398 (70°47.5´n 27°00.0´w; map 4). snow domes on milne land, supposedly resembling white pillows or cushions (= puderne). named by the geodætisk institute in 1963. púkitsivakajik – see pukkitsivakajik. púkitsivakajîp akìnarteqitâ, púkitsivakajîp kiámut kangertiva, púkit si va kajîp orqungmut kangertiva – see pukkitsivakajiip akinnarte qitaa, pukkitsivakajiip kiammut kangertiva, pukkitsivakajiip oq qum mut kangertiva. pukkelen 71ø-432 (71°08.2´n 29°16.8´w; map 4). nunatak on the west side of vindue gletscher. named by peter homewood during the 1967–72 ggu scoresby sund expeditions for its humped shape (pukkel = hump). pukkitsivakajiip akinnarteqitaa [kap dundee] 69ø-60 (69°45.3´n 23°13.0´w). penisula between manby halvø and turner ø, on the northern blosseville kyst. one of the names recorded by the geodætisk institute 1955 survey, the name derives from its location relative to pukkitsivakajik [manby halvø]. (púkitsivakajîp akìnarteqitâ.) pukkitsivakajiip kiammut kangertiva 69ø-57 (69°52´n 23°16´w). fjord on the northern blosseville kyst. the name was recorded by the geodætisk institute 1955 survey, and translates as ‘the fjord with pukkitsiakajik to its north’. (púkitsivakajîp kiámut kangertiva). pukkitsivakajiip oqqummut kangertiva [deichmann fjord] 69ø-22 (69°49.0´n 23°14.0´w). fjord sw of pukkitsivakajik [man by halvø]. one of the names recorded by the geodætisk institute 1955 survey, the name translates as ‘the fjord in the lea of pukkitsivakajik’. (púkitsivakajîp orqungmut kangertiva.) pukkitsivakajik [manby halvø] 69ø-5 (69°49.0´n 23°04.0´w). peninsula on the north blosseville kyst, sw of kap brewster. the name was recorded by the geodætisk institute 1955 survey, and translates roughly as ‘the little low’, a reference to its relative prominence. (púkitsivakajik, pukkitsukajik.) pukugkiarpik – see pukukkiarpik. pukukkiarpik 70ø-289 (70°30.0´n 22°15.5´w). hillside ne of ittaa jimmit [kap hope], sw liverpool land. recorded by the 1955 geodætisk institut name registration, the name means ‘where one picks berries to take home’. (pukugkiarpik.) pulfrichfjellet 74ø (74°22.3´n 21°13.7´w). mountain ridge on north clavering ø. used only on nsiu maps (lacmann 1937), and named after carl pulfrich [1858–1927], a german scientist who was one of the founders of the photogrammetric developments of carl zeiss, jena. 282 punktum 76ø-137 (c. 76°22´n 26°52´w; fig. 21). small nunatak 2175 m high in sw dronning louise land, so named by j.p. koch’s 1912–13 expedition because it was the last nunatak passed before crossing the inland ice (punktum = full stop). punta celso gilberti 70ø (70°04.8´n 23°01.0´w). mountain 1262 m high west of milano gletscher on volquaart boon kyst. it was climbed by leonardo bonzi’s 1934 expedition, and named after an italian climber killed in a mountaineering accident in 1933. (gil berti peak.) punta club alpino italiano 70ø (70°03.0´n 22°32.2´w). mountain on the volquaart boon kyst, the present sfinxen. it was climbed by leonardo bonzi’s 1934 expedition. (p. cai, club alpino italiano peak.) punta karfen 72ø (c. 72°08´n 24°58´w). peak in the vikingbræ region, climbed by g. dionisi’s 1982 expedition. punta roma 70ø (70°03.8´n 22°51.5´w). mountain 1267 m high west of roma gletscher on volquaart boon kyst, the present bul bjerg. it was climbed by leonardo bonzi’s 1934 expedition. (rome hill, rome peak.) punta umberto balestreri 70ø (70°03.4´n 23°08.7´w). mountain 1636 m high on volquaart boon kyst, the present isjomfruen. it was climbed by leonardo bonzi’s 1934 expedition, and dedicated to the president of the club alpino accademico who had died in a mountaineering accident in 1933. (p. balestrieri, balestreri peak.) punta degli italiani 70ø (70°01.1´n 22°58.8´w). mountain 1701 m high on volquaart boon kyst, the present pyramiden. this was the highest peak climbed by leonardo bonzi’s 1934 expedition. (p. italiani, peak of the italians.) purpurfjeld 72ø-137 (72°10.9´n 22°27.7´w). mountain on south traill ø, se of drømmerbugt. the name came into use during lauge koch’s geological expeditions in the 1930s, and is attributed to helge backlund. it derives from the colour of the rocks. purtscheller tinde 71ø (71°52.6´n 25°37.0´w). mountain between spærregletscher and prinsessegletscher. first climbed by the 1967 berchtesgaden expedition, and named probably after l. purt scheller who made notable climbs in the alps and on kilimanjaro in the 1870s and 1880s. puslingen 72ø (72°40.2´n 22°33.3´w). small island in vega sund, west of nordenskiöld ø. used on the nsiu maps of lacmann (1937), and named for the shape (pusling = goblin, tiny tot). pusterdal 76ø-255 (76°54.5´n 21°14.3´w; map 4). valley at the head of pustervig. the name was first used as a reference locality in the meteorological reports of the 1906–08 danmark-ekspedi tionen. (pustertal, pustervigdalen.) pusterelv 76ø (76°55.0´n 21°07.3´w). name very occasionally used for the river draining into pustervig. it appears on a map by charles s. poulsen, youngest member of the the 1906–08 danmark-eks peditionen, which was reproduced by lundbye (1984). pustersø 76ø (76°54.2´n 26°16.3´w). name used on a sketch map by charles s. poulsen reproduced by lundbye (1984), for a lake in pusterdal. see also pusterelv. pustervig 76ø-146 (76°55.1´n 21°00.0´w). deep small bay on the south side of mørkefjord, south of danmarks monumentet. named by the 1906–08 danmark-ekspeditionen. it was originally known as lysevig. pustervig 76ø (76°55.3´n 21°01.6´w). hut used as a meteorological station by peter freuchen from september 1907 to april 1908, erected on the west side of pustervig by the 1906–08 danmarkekspeditionen. only the stone walls reinforcing the hut remain; the wood was removed for use as firewood by hvalrosodden hunting station in 1920 (p.s. mikkelsen 1994). it was also known as freuchens hytte. (puster-wigh, puster cove, byljavik.) pututaajik 70ø-223 (70°42.3´n 21°39.4´w). small bay on the north side of the heywood bjerge, south liverpool land. recorded by the 1955 geodætisk institut name registration, the name means ‘it has a hole’, and refers to a natural hole or cave in the cliff. (pututâ jik.) pututâjik – see pututaajiik. pyramid peak 71ø (71°59.3´n 25°27.5´w). prominent rock peak 2293 m high south of dammen, central stauning alper. so named and first climbed by the 1964 aac zürich expedition. published geodetic institute maps apply the name queens tinde to this peak. pyramide 72ø (72°05.9´n 25°42.7´w; map 5). this 2250 m high mountain is described as the dominant peak at the head of the tre kantgletscher basin. climbed and so named by wolfgang wein zierl’s 1970 expedition. location a little uncertain. pyramidedalen 76ø-222 (76°57.9´n 20°37.1´w). valley west of fuglenæbsfjeldet on the north side of mørkefjord. named by eigil knuth’s 1938–39 mørkefjord expedition for the cones of moraine in the valley. pyramideelv 80ø-108 (80°26.5´n 21°00.0´w). river draining ne from keglen into the head of ingolf fjord. probably named by john haller following explorations during lauge koch’s 1956–58 expeditions. pyramidefjeld 72ø-254 (72°08.5´n 24°52.6´w). mountain between the head of vikingebræ and bersærkerbræ, north of hjørnespids. the name is found first in the form pyramide in a report by braun (1953), who made an unsuccessful attempt from the north side while assisting erdhardt fränkl during lauge koch’s 1951 expedition. it was named for the shape as seen from the north. in mountaineering literature it generally goes under the name kensington. pyramiden 70ø-276 (70°01.1´n 22°58.8´w). prominent pyramidshaped peak 1701 m high on volquaart boon kyst, so named during the 1931–34 treårsekspeditionen by laurits bruhn for its shape. it has also been called punta degli italiani. pyramiden 74ø-223 (74°01.3´n 21°34.2´w). feature on the north slope of frebold bjerg, nw hold with hope, between river 8 and river 9. named during the 1931–34 treårsekspeditionen by eigil nielsen for its shape. pytelv 70ø-164 (70°46.6´n 22°25.7´w). small river in south liverpool land draining west into hurry inlet, so named during the 1931–34 treårsekspeditionen by laurits bruhn because of its small size (pyt = puddle). pythagoras bjerg 71ø-69 (71°22.9´n 25°14.4´w). mountain south of holger danske briller on the north side of nordvestfjord. the name originated during the 1931–34 treårsekspeditionen and was adopted at the suggestion of r. spärck. in shape the mountain ap proximates to a right-angled triangle. the name first appeared on the maps of thorson (1934). (pythagoras-gebirge, mt. pythagoras.) pytten 76ø-228 (76°57.1´n 22°00.0´w). westernmost lake in vig fus dal, west of mørkefjord. the name was given by the place name committee in 1940 to replace a suggestion by eigil knuth’s 1938–39 mørkefjord expedition (pyt = puddle). pælen 70ø-37 (70°45.0´n 25°19.9´w). hill in east milne land between charcot havn and kap leslie, named during the 1931–34 treårsekspeditionen by hermann aldinger as pfahl (pæl = pole, stake). pøllen 71ø-120 (71°18.9´n 21°46.1´w). peninsula in east liverpool land, north of trekanten. the name was proposed by the place name committee in 1938, and was named for its shape (pøllen = the cushion). it was originally thought to be an island. pølseneset 72ø (72°55.1´n 21°55.2´w). narrow peninsula on east geographical society ø, nw of kap mackenzie. so named on the nsiu maps of lacmann (1937) for the sausage-like shape. (pölse neset.) pølsevika 72ø (72°55.0´n 21°57.5´w). small bay on east geogra ph ical society ø, nw of kap mackenzie near pølseneset. used on the nsiu maps of lacmann (1937). (pölsevika). påskedalen 76ø-190 (76°08.1´n 19°56.6´w; map 4). valley in ad. s. jensen land between syttendemajfjorden and påskenæsset. pro posed by nanok, the name first appeared on a map in jennov (1939). 283 påskehytten 76ø-210 (76°09.8´n 19°47.6´w). danish hunting hut about 2 km south of påskenæsset on the east coast of ad. s. jensen land, built by nanok in august 1938. (paaskehytten, påskenæs hytten, påskenæsset.) påskenæsset 76ø-13 (76°09.7´n 19°47.2´w; map 4). peninsula on the east coast of ad. s. jensen land, named by the 1906–08 dan mark-ekspeditionen in the form paaskenæsset. henning bistrup and håkon jarner visited the area at easter 1908 (j. løve, personal communication 2009). (easter naze, paaskenæs.) q qaaliartalik 70ø-247 (70°56.8´n 21°38.1´w). peninsula on the east coast of liverpool land, of which kap greg is the east cape. one of the names recorded by the 1955 geodætisk institut name registration, it translates as ‘where there is a hole’. (qâliartalik.) qagtiterpâjik – see qattiterpaajik. qâliartalik – see qaaliartalik. qámavai, qámavâjivata ilertâ – see qammavai, qammavaajivata ilerta. qammavaajivata ilerta 70ø-364 (70°29.7´n 21°58.2´w). small bay west of scoresbysund, part of nw hvalrosbugt. recorded by the 1955 geodætisk institut name registration, it translates as ‘hunting place bay’, referring to its proximity to palasip qammavaajua. (qá ma vâ jivata ilertâ.) qammavai 70ø-315 (70°28.1´n 21°56.9´w). peninsula on the east side of rosenvinge bugt, south liverpool land. recorded by the 1955 geodætisk institut name registration, the name means ‘where one lies in wait hunting’. (qámavai.) qaqilaasivik 71ø-207 (71°18.5´n 25°08.8´w). point west of sydkap near the mouth of nordvestfjord. recorded by the 1955 geodætisk institut name registration, the name means ‘the depot place’. it was a location where seals were taken up onto land and covered with stones, and later collected. (qaqilausivik.) qaqilaasivik kangitteq 71ø-206 (71°20.8´n 25°13.7´w). place on the coast west of sydkap, at the foot of the mountain pythagoras bjerg. recorded by the 1955 geodætisk institut name registration, the name translates as ‘the western depot place’. (qaqilausivik kangigteq.) qaqilausivik, qaqilausivik kangigteq – see qaqilaasivik, qagilaasivik kangitteq. qaqqaqqaap inaa 70ø-318 (70°27.2´n 21°57.0´w). small lake on the east side of rosenvinge bugt. recorded by the 1955 geodætisk institut name registration, the name translates as ‘the home of the diver’, and refers to the red-throated diver that nests in small lakes. (qarqarqâp unâ.) qarqarqâp unâ – see qaqqaqqaap inaa. qarrtsilunisletta 72ø (72°43.5´n 22°11´w; fig. 14). south-facing lower slopes of se geographical society ø, nw of kap mc clintock. so named on nsiu maps of lacmann (1937) after the lockheed-vega aeroplane ‘qarrtsiluni’ loaned to nsiu by consul lars christensensen for the 1932 aerial photography carried out in east greenland. the name apparently can be translated as ‘soul of the whale’, but in alaskan inuit it means artistic concentration, literally a stillness waiting for something to break. qassuserpaajik – see qattiterpaajik. qattiterpaajik [fox pynt] 70ø-313 (70°28.2´n 21°56.7´w). peninsula on the east side of rosenvinge bugt, south of scores bysund, south liverpool land. one of the names recorded by the 1955 geodætisk institut name registration, it translates as ‘the place where one puts out nets’. the scoresbysund local newspaper recorded in 1984 the spelling qassuserpaajik. (qagtiterpâjik.) qeqertaq 71ø-76 (71°48.8´n 22°45.5´w). island on the north side of fleming fjord, named during the 1931–34 treårsekspeditionen by arne noe-nygaard in the form kekertak, the greenlandic word for an island. qeqertaq prins henrik 77ø-1a (77°43.0´n 17°45.0´w). on 11 june 2004, the greenland home rule authority (hjemmestyret) officially changed the name of the island île de france to qeqertaq prins henrik as a present to prince henrik of denmark, husband of queen margrethe ii of denmark, on the occasion of his 70th birthday. qernerterajik 70ø-339 (70°27.3´n 21°37.8´w). cape between kap swainson and kap lister, south liverpool land. one of the names recorded by the 1955 geodætisk institut name registration, it means ‘the blackish’, and refers to its dark colour. the local scores bysund newspaper recorded in 1984 the spelling qernerdaajik. qilalugkat nunât – see qilalukkat nunaat. qilalukkat nunaat 71ø-214 (71°15.1´n 25°27.6´w). coastal stretch of east renland, north of skillebugt. the name was recorded by the 1955 geodætisk institut name registration, and means ‘white-fish land’. this name probaly refers to attempts to locate halibut banks and establish a fishery in the region in the 1940s. (qilalugkat nunât.) qilernaq 71ø (71°34.5´n 23°57.7´w). locality where the rivers from coloradodal and major paars dal meet in ørsted dal, where colo ra do dal hytten was built in 1983. the name means ‘the place where something meets’. qíngâjiva, qíngajivata qáqartivâ – see qinngaajiva, qinngajivata qaqqartivaa. qinngaajiva 70ø-303 (70°30.1´n 22°00.3´w). ne side of hvalros bugt, south liverpool land. one of the names recorded by the 1955 geodætisk institut name registration, the name means ‘its little bottom’. (qíngâjiva.) qinngajivata qaqqartivaa [mågefjeld] 70ø-199 (70°31.6´n 21°54.0´w). hill north of scoresbysund town, south liverpool land. the name was recorded by the 1955 geodætisk institut name registration, and translates as ‘the bay’s big hill’. the scoresbysund newspaper recorded in 1984 the local name nerdiit iaat. (qíngaji vata qáqartivâ.) qooroq 71ø-78 (71°40.9´n 23°15.1´w). valley on the west side of the the head of fleming fjord. named by a. noe-nygaard during the 1931–34 treårsekspeditionen in the form quôrok, the green landic word for a valley. (qôroq.) queens tinde 71ø-365 (71°59.3´n 25°27.5´w; map 5). snow peak 2293 m high on the sw side of krabbe gletscher. climbed by 1963 cambridge university expedition on 30 july, and named after queens’ college, cambridge. one of the most picturesque of cam bridge colleges, it was founded in 1448 by the wife of henry vi, and refounded in 1465 by the wife of edward vi. it has also been known as pyramid peak. the name is slightly misplaced on published geo detic institute maps. (queenstinde.) quellpingo 72ø (72°33.6´n 23°33.4´w). name used by fritz müller during lauge koch’s 1954–55 expeditions after a pingo beside ka rup elv, traill ø. it is the smallest of a group of three pingos, 9 m high, and contains a spring (= quell) flowing at the rate of 1.4 litres a second (müller 1959). quensel bjerg 71ø-86 (71°36.6´n 22°44.8´w). mountain on wegener halvø. named during the 1931–34 treårsekspeditionen by arne noe-nygaard as mt. quensel, after percy dudgeon quen sel [1881–1966]. a swedish igneous petrologist he was professor at the university of stockholm, and noted for his work on charnockites and syenites. (quenzels bjerg.) quest-hytten – see schlelderup-hytta. qúpaulakajik, qúpaulakajik kangigteq – see quppaalakajik, qup paa la kajik kangitteq. qúpaulartivakajik – see quppaalartivakajik. quppaalakajik [rendeelv] 70ø-184 (70°32.3´n 22°22.8´w). river and ravine in south liverpool land draining into hurry inlet. re corded by the 1955 geodætisk institut name registration, the name translates as ‘the little ravine’. (qúpaulakajik.) quppaalakajik 70ø-147 (70°30.7´n 22°37.1´w). ravine in neill klinter on the west side of hurry inlet. the name was recorded by 284 the 1955 geodætisk institut name registration, and means ‘the little ravine’. dinosaurus klöft has also been used. (qúpaulakajik.) quppaalakajik kangitteq [hulelv] 70ø-183 (70°33.4´n 22°24.4´w). river and ravine north of quppaalakajik [rendeelv], draining into hurry inlet. one of the names recorded by the 1955 geodætisk in stitut name registration, it translates as ‘the outer little ravine’. (qúpaulakajik kangigteq.) quppaalartivakajik 70ø-203 (70°30.6´n 21°33.3´w). ravine in se liverpool land near kap lister. recorded by the 1955 geodætisk institut name registration, the name translates as ‘the big ravine’. the scoresbysund newspaper recorded in 1984 the local spelling qappaalardivakajik. (qúpaulartivakajik.) r raatiuup nuaa 70ø-368 (70°29.0´n 21°57.4´w). cape at the mouth of elvdal in scoresbysund. the name was recorded by the 1955 geodætisk institut name registration, and translates as ‘radio cape’, referring to the former radio station on top of the low hill behind the cape. (râtiûp nûa.) raatiuup tunua 70ø-371 (70°29.1´n 21°56.4´w). slope east of scoresbysund, east of the former radio station. recorded by the 1955 geodætisk institut name registration, the name means ‘the radio station’s back side’. in 1927–28 janus sørensen built a radio station and seismic station on top of the 60 m high hill. (râtiûp tunua.) rabbit ears island 78ø (78°00.0´n 18°52.6´w). large island in the danske øer group, named by jane a. gilotti for its shape. the name is used as a geological reference locality in reports of studies on ultrahigh pressure eclogites (lang & gilotti 2007). rabsontinde 72ø (72°00.7´n 25°10.0´w; map 5). minor outlying peak of kapelle, about 1640 m high, on the north side of sefström gletscher. climbed by the 1998 scottish mountaineering club expedition. (robson.) raceway 71ø (c. 71°25´n 22°33´w). locality in jameson land where farish a. jenkins during the 1988–89 harvard university palaeontological expeditions discovered spectacular fossil dinosaur tracks. a network of 52 different tracks are preserved. radiobæk 76ø-294 (76°55.8´n 20°19.8´w). stream between mørke fjord station, which was also a radio station, and termografengen. so named by the 1938–39 mørkefjord expedition. raffles ø [appalik] 70ø-209 (70°36.1´n 21°31.2´w; map 4). is land off the coast of south liverpool land. named by william scores by jr. in 1822 as raffles island out of respect to the revd tho mas raffles [1788–1863], a prominent independant minister, who held the living at great george street, liverpool from 1812 to 1862. (île raffles, raffles ö, raffle ø, rafle insel.) raffles sø 70ø (70°35.6´n 21°32.4´w). lake on raffles ø where ma terial was collected for radiocarbon age determinations (cremer et al. 2008). ragekniven 76ø (76°20.5´n 20°23.6´w). name used for a mountain on nanok ø, about 5 km sw of teufelkap in roon bugt. accor d ing to poulsen (1991) the name was given by håkon jarner (rage kniv = razor). ragnhildshytta 73ø (73°28.5´n 25°02.9´w). norwegian hunting hut at the east side of the mouth of grejsdalen, east andrée land. built in march 1937 for arktisk næringsdrift by magne råum, and named after råum’s girlfriend, ragnild lien. the hut has also been known as grejsdalshytten and eleonorebukta. (ragnilds hytte.) ragnas ø 76ø (76°27.0´n 20°54.5´w). name used by henning bi strup during the 1906–08 danmark-ekspeditionen for the present godfred hansen ø. it was probably given for one of henning bistrup’s family (j. løve, personal communication 2009). olgas ø has also been used. ramnenuten – see ravnebjerg. ramnfjellet 73ø (73°25.0´n 23°09.1´w). mountain on the south side of gauss halvø, corresponding to the present wiman bjerg. so named on an nsiu map (1932a), and derived from the norwegian dialect word for a raven. there are many similar place names in norway. ramp 77ø (c.77°13´n 24°00´w). upper part of britannia gletscher, dronning louise land. the name was given by members of the 1952–54 british north greenland expedition (banks 1957) after the notable ‘ramp’ which features in accounts of scott’s 1910–13 antarctic expedition. rampe 71ø-319 (71°37.7´n 24°26.1´w). sand bank on the west side of schuchert flod river bed, at the front of bjørnbo gletscher. named by enrico kempter during lauge koch’s 1956–58 expeditions. rampevæggen 70ø-424 (70°37.6´n 28°49.0´w). mountain wall north of rolige bræ forming a boundary to a ramp-like tongue of ice. named by laurent jemelin during the 1967–72 ggu scoresby sund expeditions. ramsay bjerg 73ø-71 (73°30.5´n 22°42.7´w). mountain on gauss halvø. named during lauge koch’s 1929 expedition in the form mt. ramsay by helge g. backlund after wilhelm ramsay [1865– 1928], a finnish mineralogist and petrologist and an authority on fennoscandian structures. he was professor of geology at the uni versity of helsingfors. ran øer 72ø-328 (72°17.2´n 23°54.2´w). small islands or skerries on the sw side of kong oscar fjord. named by a.g. nathorst in 1899 as rans skär, according to his narrative after ‘den falska ran’; this features in the swedish masterpiece, tegnér’s fritiofs saga. in 1957 the authorised spelling was changed from rans skær to ran øer at the suggestion of søkortarkivet (nautical charts archive), who considered the islands were too large to be called skerries. (rans øer, rans rock.) randbjerg 73ø-665 (73°41.0´n 25°26.3´w). ridge between ende løs and spaltegletscher, andrée land. named during lauge koch’s 1948–50 expeditions by erdhardt fränkl (rand = edge). randbøldalen 73ø-346 (73°20.3´n 22°14.5´w). valley in the south giescke bjerge, draining eastwards. the name was proposed by the place name committee in 1939 after the area of randbøl in midjylland where there are more than 300 burial mounds. björndalen and franklindalen have also been used. randelven 76ø (76°54.8´n 22°04.6´w). name used by j.p. koch’s 1912–13 expedition for a large river at the margin of the inland ice in west daniel bruun land, flowing into borgfjorden (rand = mar gin). randen 74ø-328 (74°09.0´n 24°06.0´w). high plateau in ole rømer land west of vibeke sø. named during lauge koch’s 1936– 38 expeditions by heinrich bütler, probably for the locality of the same name near schaffhausen. randenæs 72ø-140 (72°25.5´n 25°43.0´w). peninsula on the north side of forsblads fjord. so named by eugéne wegmann during the 1931–34 treårsekspeditionen, probably because the peninsula marks an important geological boundary. randers fjord 70ø-251 (70°57.8´n 21°46.7´w; map 4). fjord on the east coast of central liverpool land, named during the 1931– 34 treårsekspeditionen by laurits bruhn after the fjord of the same name on the east coast of jylland, denmark. randspids 71ø-290 (71°51.9´n 24°08.0´w; map 5). mountain at the sw margin of the werner bjerge (rand = margin). named by peter bearth and eduard wenk during lauge koch’s 1953–54 expeditions and climbed by bearth in 1954. randsøen 76ø-124 (76°41´n 22°58´w; map 4). lake at the west margin of storstrømmen, so named by j.p. koch’s 1912–13 expedition (rand = margin). (randsee.) ransey 75ø, 76ø (c. 76°15´n 18°42´w). name given to an island on the north part of the east coast of greenland on a 1706 map by torfæus. according to tornøe (1944), ransey might correspond to the present store koldewey. (drangey.) 285 ranunkeldal 71ø-444 (71°30.5´n 24°08.6´w). valley on the east side of schuchert dal, connecting with the upper reaches of ørsted dal. named by geoffrey halliday following botanical work during the 1971 northern universities expedition, for the occurrence of a rare buttercup (ranunculus pedatifidus). ranvik 74ø (74°26.9´n 20°24.8´w). small bay on the sw coast of wollaston forland, east of zackenberg bugt. used only on nsiu maps (lacmann 1937), and named after the home of lars chri stensen in sandefjord, norway. see also lars christensenfonna. rasmussen spids 71ø (71°50.0´n 25°37.2´w; map 5). mountain 2468 m high in the ne part of the borgbjerg gletscher region, southern stauning alper. probably named by the 1977 schwä bische stauning alper expedition. râtiûp nûa, râtiûp tunua – see raatiuup nuaa, raatiuup tunua. rath boon insel – see rathbone ø. rathbone ø [immikkeertikajik] 70ø-221 (70°40.3´n 21°28.0´w; maps 3, 4). island off the coast of south liverpool land with a peak resembling the ruins of a castle. it was named by william scoresby jr. in 1822 as rathbone island (fig. 3) after an esteemed friend, william rathbone [1787–1868], who in 1837 became mayor of liverpool. scoresby went on holiday to ireland with rathbone and thomas traill in 1820. a party led by helge g. backlund climbed to the summit in june 1933. bobé (1936 p. 45) suggested that volquaart boon, who was swept into the mouth of scoresby sund in 1761, had given an island the name rath boon insel, adopted by scoresby (1823) as rathbone island. the idea appears to derive from a map drawn by boon and at one time owned by m. worm skiold, but subsequently lost in a fire. the place name committee considered the problem in 1960–61, and concluded the story im probable. (rathbones ö, île rathbone.) rattenfanger peak 73ø (73°32.0´n 26°09.5´w). snow peak 2155 m high on the south side of grejsdalen, andrée land. climbed by the 2007 army boreal zenith expedition. räuberloch 75ø (c. 75°19´n 17°48´w). feature in the vicinity of the base camp of the 1943–44 operation bassgeiger at kap sussi, shan non. the name is recorded by olsen (1965), but the exact location is uncertain. raudalen 73ø (73°05.9´n 23°56.7´w). norwegian hunters name for barnabas dal on south ymer ø. possibly named after the valley of the same name in the jotunheimen, or several other areas in nor way. raudalshytta 73ø (73°04.6´n 23°43.3´w). norwegian hunting hut on the north side of sofia sund east of the mouth of barnabas dal (which the norwegians called raudalen), built by arktisk næ rings drift. the original hut was moved to renbugten in august 1931, and replaced in august 1932 by the former strømhytta. the names stor-dalen, bødtker, dalhytten and barnabasdal hytte have also been used. (røvdallen.) raudeberg 72ø (72°24.4´n 24°54.6´w). norwegian hunting hut on the south coast of segelsällskapet fjord 1 km sw of skipperdal, built in 1930 by the møre expedition. the area had no apparent name, and was called raudeberg by the hunters for the redcoloured rocks. (raudberget.) raukelv 70ø-101a (70°28.3´n 23°11.0´w). river in south jameson land. the name was first used in the form rauk river by aldinger (1935), and approved following the 1967–72 ggu scoresby sund expeditions. it is derived from a scandinavian word for a certain kind of rock formation in gotland. raukplateau 70ø-402 (70°29.8´n 23°12.0´w). low plateau in south jameson land beside raukelv. the name was first used in a report by aldinger (1935) in the form rauk plateau, but was not approved until 1972. see also raukelv. raven glacier 72ø (72°00.7´n 24°47.5´w). minor glacier on the east side of storgletscher, named by the by the 2007 smc east green land expedition for a sighting of a raven. raven pond 76ø (76°16.6´n 18°36.3´w). small lake on store kolde wey where sampling was undertaken for phytoplankton studies (cremer et al. 2005). ravin du solitaire 70ø (c. 70°51´n 22°23´w). ravine on the west coast of liverpool land between kalkdal and bodal, draining west into hurry inlet. the name was used by rothé (1941) in his de scription of the french international polar year 1932–33 work at scoresbysund. ravin du suisse 70ø (70°45.5´n 22°25´w). ravine on the west coast of liverpool land, draining west into hurry inlet, probably that carrying the river damelv. the name was used by rothé (1941) in his description of the geological results of the french international polar year 1932–33 at scoresbysund. ravn pynt 76ø-156 (76°08.5´n 18°31.6´w; map 4). locality on the east coast of store koldewey. named by lauge koch’s 1926–27 expeditions as pt. ravn for jesper peter johansen ravn [1866– 1951], a danish geologist and palaeontologist, and museum in spector at the mineralogical museum, copenhagen from 1907 to 36. he had briefly described geological work carried out at this locality during the 1906–08 danmark-ekspeditionen. ravn’s ravine 76ø (76°17.9´n 18°37.2´w). ravine on the east coast of store koldewey about 2 km north of nordre gneisnæs, where eigil nielsen collected fossils in 1933. the name was used by fre bold (1935) and maync (1949). see also ravn pynt. ravnas bre 71ø (71°54.3´n 25°15.5´w; map 5). name given to a northern branch of roslin gletscher by the 1996 norwegian stau ning alper expedition. it was named after ole ravna [1841–1906], who accompanied fridtjof nansen on his crossing of the inland ice in 1888. it has also been called newnham glacier. ravnebjerg 73ø-175 (73°35.4´n 21°17.8´w; map 4). mountain in the southern tågefjeldene. the name derives from the ramne knuten of an nsiu map (1932a), but is now used in a wider sense than the nsiu usage to cover also their localities dyrhö, dyrfjellet, blåhö and gråbeinryggen. the name derives from the norwegian dialect word for a raven (= ramn). ravneknuten has also been used. ravnedal 70ø-357 (70°08.1´n 22°13.4´w). small valley ne of the settlement at kap brewster. name used by hassan (1953) in his description of material collected during lauge koch’s 1951 expedition. the name was given for the ravens. ravnedalen 76ø-271 (76°19.4´n 21°48.8´w; map 4). valley in east rechnitzer land, south of ravnefjeldet. so named by the 1938–39 mørkefjord expedition for the ravens. ravnefjeld 71ø-83 (71°42.3´n 22°41.5´w). mountain on wegener halvø, se of vimmelskaftet, so named during the 1931–34 treårs ekspeditionen by arne noe-nygaard for the ravens. ravnefjeldet 76ø-270 (76°21.5´n 21°45.0´w; map 4). mountain in east rechnitzer land. so named by the 1938–39 mørkefjord expedition, for the presence of ravens. ravnefjellet 71ø (c. 71°46´n 22°26´w). mountain near kaares-bu, possibly on wegener halvø south of kap brown. the name was used by the norwegian hunters helge ingstad and normann and er sen in 1932, because during their first night in their first camp they saw a raven flying in the direction of this twisted peak. ravnekløft 70ø-321 (70°25.1´n 21°58.6´w). gulley near kap tobin in south liverpool land. named during the 1924–25 colonisation expedition for the ravens, which were seen in flocks of 9–12 here (pedersen 1926). (ravnekløften, raven cleft.) ravnenæs 70ø-252 (70°59.0´n 21°46.0´w). peninsula between maria ger fjord and randers fjord, liverpool land. so named by lau rits bruhn during the 1931–34 treårsekspeditionen, for the ravens. ravnestenen 74ø (74°28.7´n 20°34.3´w). reference locality used by visiting scientists to zackenberg forskningsstation. rebild 72ø-91 72ø-92 (72°48.1´n 23°59.0´w; map 4). name used for the mountain range in ne traill ø, and also originally for the valley to its south (located at 72°46.1´n 24°06.9´w) which was later renamed de lemos dal. the names have often been used in 286 the forms rebild bjerge and rebild dal. named by ove simonsen during the 1931–34 treårsekspeditionen after the danish locality rebild south of ålborg, jylland. rechnitzer land 76ø-187 (76°19.0´n 22°00.0´w; maps 2, 4). land area between soranerbræen and bræfjord. mapped by lauge koch during flights in 1932 on the 1931–34 treårsekspeditionen, it was named after vice-admiral hjalmar rechnitzer [1872–1953], who was director of the marine ministry from 1923 to 32 and head of søværnskommandoen from 1932 to 40. (rechnitzerland.) rechnitzerhytten 76ø-206 (76°20.2´n 21°49.8´w). danish hunting hut on the east coast of rechnitzer land, built by nanok in august 1938. now a ruin. see also rechnitzer land. red rose mountain 70ø (70°51.7´n 26°05.7´w). summit 2067 m high on the north side of korridoren, milne land. climbed by the 2004 west lancashire scouts expedition during a ski tour. redekammen 76ø-150 (76°56.1´n 21°28.4´w). mountain ridge on the south side of inner mørkefjord with many minor peaks. named by the 1938–39 mørkefjord expedition presumably for its appearance (redekam = comb). soel-backen has also been used. regnbuedal 71ø-344 (71°53.8´n 22°49.3´w). valley draining into the north side of fleming fjord at the mouth of ørsted dal. so named during lauge koch’s 1958 expedition by k. grasmück and rudolf trümpy, on account of the vivid colours of the triassic strata forming its flanks. regnbueklippe 76ø-313 (76°59.0´n 24°39.5´w; map 4). cliff on the south side of admiralty gletscher, dronning louise land. named by the 1952–54 british north greenland expedition as regn bue klippe because the coloured rock units were reminiscent in colour and shape of a rainbow. regnelv 71ø-189 (71°18.1´n 24°26.3´w). river in west jameson land draining sw to enter the sea north of gurreholm. so named during the 1936–38 two-year expedition by hans stauber, because of his experiences here during the rainy summer of 1937. regntoppen 71ø-395 (71°40.0´n 22°51.8´w). mountain 810 m high on wegener halvø. so named by katharina perch-nielsen during the 1967–72 ggu scoresby sund expeditions because it rained while camped here after several weeks of sunshine. regolitplateau 71ø-133 (71°04.5´n 22°10.6´w). plateau area on the north side of inner storefjord, central liverpool land. so named by helge g. backlund during the 1931–34 treårs eks peditionen, for the regolith, the frost-shattered bedrock which covers the plateau areas. reiat 74ø-327 (74°08.6´n 23°36.3´w). high plateau on the north side of vibeke sø. named by heinrich bütler during lauge koch’s 1936–38 expeditions, probably for the hill of the same name near schaffhausen, switzerland. reinaelv 73ø-152 (73°34.9´n 20°49.3´w). river in se hold with hope, named on an nsiu map (1932a) as reina, for the many traces of reindeer. reinhard bjerg 71ø, 72ø (71°57.7´n 28°11.0´w). name used by helge g. backlund during the 1931–34 treårsekspeditionen for the present backlund bjerg on the north side of inner nord vestfjord. reinsbukta 73ø (73°21.0´n 26°28.0´w). norwegian hunting hut on the north side of renbukten in south andrée land. the original hut was brought here in 1930 from the mouth of barnabasdal in sofia sund. it has also been known as ha-ha-hytta. (reinli, ren bugt hytten.) rejedal 70ø-359 (70°08.7´n 22°12.1´w). small valley ne of the settlement at kap brewster. name used by hassan (1953) in his de scription of fossils collected here during lauge koch’s 1951 expedition, and given for the occurrence of fossil crabs. rekdalsundet 72ø (72°41.7´n 22°28.5´w; fig. 14). sound between se geographical society ø and nordenskiöld ø. used on the nsiu maps of lacmann (1937), the name commemorates the norwegian skipper hans rekdal [b. 1899], commander of the vesle kari on the 1929 nsiu expedition. rekstadfjellet 73ø (73°58.7´n 22°19.0´w). mountain in east hudson land. so named on nsiu maps of lacmann (1937) after john bern hard rekstad [1852–1934], a norwegian geologist and glaciologist. rekvedøen 77ø-20 (77°19.8´n 18°59.3´w; map 4). island on the south side of skærfjorden. it was originally named ilot del rosio by the duke of orléans in 1905, but the 1906–08 danmark-ekspedi tionen mistakenly gave the name rosio to an island farther east, and named this island rekved-øen because of finds of driftwood. the error of position was soon discovered, but was considered not worth the confusion correction of the position would cause. (re kve dön, driftwood island.) remigolepisryg 73ø-282 (73°27.9´n 23°11.1´w). mountain ridge on gauss halvø, so named during the 1931–34 treårsekspe ditionen by gunnar säve-söderbergh (säve-söderbergh 1934) be cause of finds of several richly fossiliferous horizons containing ‘remigolepis’. (remigolepis ridge.) renbjerg 73ø-681 (73°34.1´n 26°51.0´w). mountain in west andrée land, at the head of rendal. named during lauge koch’s 1949–51 expeditions by john haller. renbugten 73ø-520 (73°20.0´n 26°28.5´w; map 4). pronounced bay on the north side of isfjord. named by a.g. nathorst’s 1899 expedition as renbukten, because a flock of 12 reindeer was seen here. this was the largest flock seen during the expedition, and the last living reindeer to be seen in east greenland. the east green land reindeer died out during the winter of 1899–1900 (reindeer bay, reinbukta, renbukta.) renbugthytten – see reinsbukta. rencontre dal 71ø-383 (71°28.0´n 29°00.0´w). major e–w-trending valley extending westwards from the head of flyverfjord as far as rencontre sø. named during the 1967–72 ggu scoresby sund expeditions after rencontre sø. rencontre sø 71ø-374 (71°29.3´n 29°20.8´w; map 4). lake at the head of rencontre dal, at the south boundary of hinks land. named by p. vogt during lauge koch’s 1957 expedition for a meeting place during field work. rendalen 73ø-641 (73°26.5´n 26°41.4´w; map 4). large valley in sw andrée land draining se into renbugt. named during the 1931–34 treårsekspeditionen by ove simonsen. rendalshytta 73ø (73°05.8´n 27°18.2´w). norwegian hunting hut at paradisdal on the east side of kjerulf fjord, nw suess land. built by bjarne and oddvar akre for arktisk næringsdrift in august 1938, and named after the rendal area of norway from which the akre brothers originate. by coincidence, there are abundant antlers and bones of the now extinct east greenland reindeer around the hut. rendeelv [quppaalakajik] 70ø-184 (70°32.3´n 22°22.8´w). river in south liverpool land draining west to hurry inlet, so named during the 1931–34 treårsekspeditionen by laurits bruhn for the shape of the valley it occupies (rende = groove). renland [tuttut nunaat] 70ø-27 71ø-40 (71°15.0´n 27°00.0´w; maps 3, 4; see also fig. 83). large land area bounded by nordvest fjord, øfjord, rypefjord and edvard bay dal. so named by carl ryder’s 1891–92 expedition because numerous reindeer (rangifer tarandus eogroenlandicus) were seen during the expedition (fig. 7). reindeer died out in east greenland about 1900. (renlandet, ren land, renntier-land.) renodde 70ø-22 (70°29.0´n 28°15.0´w; map 4). peninsula on the south side of the mouth of vestfjord. so named by carl ryder’s 1891–92 expedition because the expedition shot four reindeer here. renskæret 76ø-71 (76°40.9´n 18°30.9´w). small island south of danmark havn, so named by the 1906–08 danmark-ekspedi tionen. here, as at other localities, the ground was littered with the antlers and excrement of reindeer, although they had been extinct in the region for several years. (rendyrskæret, renskær, reindeer reef.) 287 rensund 70ø-59 (70°34.2´n 26°13.4´w; map 4). narrow sound with embayments between milne land and danmark ø. named by carl ryder’s 1891–92 expedition as ren sund because reindeer were seen here. reservatet 74ø-158 (74°11.4´n 23°14.6´w; map 4). land area between irisgletscher and wordie gletscher, a small part of ole rømer land. named during lauge koch’s 1929–30 expeditions by helge g. backlund as reservation land, originally for a slightly larger area than the present (fig. 15). it was an area in which back lund had ‘reservations’ about the geological divisions present. resoluthytten 76ø (76°56.5´n 18°10.8´w). hut built by the 1938–39 norsk–franske polarekspedisjon on the outer coast of germania land, originally called margarincentralen. restbjerg 71ø-415 (71°37.5´n 23°22.1´w). mountain 1060 m high west of the head of fleming fjord. named by katharina perchnielsen during the 1967–72 ggu scoresby sund expeditions for a small outlier (= rest) of a geological formation preserved on the summit. retrætegletscher 72ø-302 (72°01.9´n 23°56.4´w). glacier in the north werner bjerge, draining west into deltadal. so named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk, because the glacier appeared to be retreating. rev-odden 72ø (72°51.8´n 23°33.7´w). name used for the minor peninsula on the north side of vega sund, and also for the norwe gian hunting hut built here by arktisk næringsdrift in 1929 (nsiu 1932c). the hut was originally called solveigs hytta, and has also been known as kapp rygg and sverdrup hytta. revaltoppe 76ø-140 (76°39.7´n 25°42.6´w; map 4). nunataks in sw dronning louise land, west of dannebrogsfjeldene, named by j.p. koch’s 1912–13 expedition as reval-toppene or revaltoppene. reval, an old nordic name for the capital of estonia, was the site of a battle between the danes and estonians in 1219 when the danish flag, the dannebrog, is said to have dropped from the sky. (reval gip fel, reval-spitzen, revaldtoppe.) revdal 71ø-320 (71°34.2´n 24°36.3´w). valley in karstryggen draining east to schuchert flod. named by enrico kempter during lauge koch’s 1956–58 expeditions for the massive reef-building carbonates (rev = reef ). revet 74ø-122 (74°21.7´n 21°51.4´w; map 2). narrow passage between west clavering ø and the east coast of payer land. the name was used by norwegian hunters from about 1927 because it is so shallow that it can be waded at low tide (revet = shallow water place). the hunting station on the west side of the channel is also referred to as revet. the hekla was stopped by shallow water on the south side of revet in 1889. reports by other norwegian skippers that clavering ø was joined to the mainland by a sand and mud bank led to discussion of whether or not clavering ø was an island (hansen 1912). revet 74ø-269 (74°21.8´n 21°51.7´w). official danish name for the norwegian hunting station on the west side of the passage revet, west of clavering ø. the original hut on this site, tyrolerheimen, was supplemented in 1928 by a larger station known to norwegians as moskusheimen. revlerne 70ø-277 (70°01.2´n 22°51.6´w). mountain ridge up to 1210 m high adjacent to roma gletscher on volquaart boon kyst, so named during the 1931–34 treårsekspeditionen by laurits bruhn for its barred appearance (revle = bank, bar). reynolds ø [immikkeertikajiit martik] 71ø-9 (70°30.5´n 21°42.6´w; map 4; fig. 72). island off the north coast of liverpool land. named reynolds island by william scoresby jr. in 1822 in compliment to descendants of the late richard reynolds of bristol. richard reynolds [1735–1816], a quaker philantropist, who retired from business a rich man in 1789, settled in bristol in 1804 and devoted himself to dispensing charity on a large scale. (reynold ö, reynold ø.) rhaetelv valley – see rhætelv. rhedin fjord 72ø-404 (72°40.0´n 26°20.0´w; map 4; fig. 52). n–s-trending fjord between gletscherland and lyell land. named by a.g. nathorst’s 1899 expedition after a swedish businessman, martin werner rhedin [1865–1930] of ellesbo, who contributed 5000 swedish kronor to the expedition. (rhedins fjord.) rhodesia peak 70ø (70°47.9´n 26°02.1´w). peak 1440 m high on the south side of korridoren, milneland, climbed by the 2004 west lan cashire scouts expedition. rhætelv 71ø-179 (71°38.0´n 23°14.0´w). river west of the head of fleming fjord. so named by hans stauber during lauge koch’s 1936–38 expeditions because the river drains through a valley formed in rhaetic rocks. rhaetelv valley is occasionally used for the valley in english publications (e.g. hall 1964). richardpynt 72ø-271 (72°53.6´n 24°47.1´w; map 4). minor cape on ne ella ø. named by john cowie during lauge koch’s 1950 expedition after his assistant, richard nielsen of copenhagen. richmond 72ø (72°04.7´n 24°29.8´w; map 5). rock peak 1650 m high on the east side of kishmul gletscher, stauning alper. first climbed by the 1963 imperial college expedition and named after the london borough, now richmond-upon-thames. richter-hytta 72ø (72°42.0´n 22°18.0´w). norwegian hunting hut on sw geographical society ø, about 4–8 km nw of kap mc clintock, at a bay norwegians called thorolf vogts bukta. built by arktisk næringsdrift in september 1929, and named after søren richter [1903–70], who helped to build it. søren richter wintered in east greenland from 1929 to 1931 and from 1935 to 1936, and led his own hunting expeditions to east greenland in 1937–1938 and 1939–1940. he spent the war years on jan mayen, and from 1946 to 1970 was norsk polarinstitut librarian. fig. 72. looking south from murray ø to rey nolds ø, kap godfred hansen and the jagged summits of liverpool land. liverpool landkap godfred hansen reynolds ø 288 richterfjellet 74ø (74°21.5´n 21°18.5´w). mountain on nw clav ering ø. used only on nsiu maps (lacmann 1937), the name was given for hans richter [b. 1897], a german surveyor who led the stereographic work on the detailed topographical map of jordan hill, and søren richter [1903–1970], a norwegian archeologist and hunter. see also richter-hytta. ridderborgen 73ø-531 (73°05.8´n 27°28.5´w). mountain 1885 m high on the west side of the mouth of kjerulf fjord. the summit resembled a ruined castle, and was named by a.g. nathorst’s 1899 expedition as riddarborgen (= baronial castle). (riddarborg.) ridderdal 73ø-532 (73°04.3´n 27°28.9´w). valley south of ridder borgen, north goodenough land. the valley was used by j.m. wordie’s 1929 expedition as a route on his ascent of petermann bjerg, and the name appears on his maps in the form riddar valley. riddergletscher 73ø-556 (73°04.1´n 27°36.0´w). glacier at the head of ridderdal, north goodenough land, named by j.m. wordie’s 1929 expedition as riddar glacier. ridge 1–12 73ø (73°58.0´n 21°19.5´w). series of minor ridges on the ne slope of stensiö plateau, designated in this form for reference purposes during the 1931–34 treårsekspeditionen. rigi 74ø (74°38.0´n 20°42.4´w). small isolated summit in nw wol la ston forland. the name was used by wolf maync (1947) in his description of work during lauge koch’s 1936–38 expeditions, because of a resemblance to the rocks of rigi, a noted viewpoint overlooking vierwaldstättersee, switzerland. rigi nunatak 72ø-445 (72°42.0´n 27°54.5´w; map 4). nunatak on the south side of the upper reaches of hisinger gletscher. so named by eugéne wegmann during the 1931–1934 treårsekspe ditionen after the swiss locality – see also rigi. wegmann visited the nunatak in august 1934. rigny bjerg 69ø-16 (69°03.0´n 26°49.0´w; fig. 73). prominent mountain 2783 m high west of the blosseville kyst, named by jules de blosseville in 1833 as mont rigny. it was probably given for the french vice-admiral henri-marie-daniel gaultier, count de rigny [1782–1835] (j. løve personal communication 2009). the mountain was identified from blosseville’s map by the 1879 ingolf expedition and its position approximately determined by g.c. amdrup’s 1898–1900 expedition. the position and altitude are incorrect on ams maps published in 1952. hauge andersson made bearings on the summit in 1967 and 1972 when surveying on the blosseville kyst for the geodetic institute, but was unable to fix its position. the mountain probably corresponds to the bláserk of the icelandic sagas, and henry hudson’s mount of gods mercie. many expeditions have explored the rigny bjerg region, and many summits have been climbed. two members of the 1998 rigny bjerg expedition climbed to a height of c. 2600 m on the steep and narrow ne ridge, before retreating due to dangerously loose snow. the first complete ascent was on 19 july 2003 by the ‘midnight sun 03 expedition. (rigny-fjæld.) riis-carstensens dyb 75ø (c. 75°41´n 18°22´w). offshore channel 400 m or more in depth between the south point of store koldewey and shannon. discovered by eigil riis-carstensen [1892–1953], a danish naval officer, when he was ice-pilot on the 1932 gefion expedition. he was also a director of nanok. rimhytta 72ø (72°43.1´n 26°10.5´w). norwegian hunting hut built east of kap hedlund in september 1934 by arktisk næringsdrift, also known as kap hedlund hytta. ole klotset gave the name when he found the hut covered in hoar-frost (= rim) in november 1934. it was replaced in 1964 by a sirius hut. (rimhytten.) rimfaxebreen 74ø (74°23.0´n 20°43.3´w). glacier on ne clavering ø. so named on nsiu maps of lacmann (1937), after rimfaxe (or rimfakse) of old nordic mythology, a horse with frost (= rim) on its mane which follows night around the world dropping morning-dew from its bit. ringøen 76ø-275 (76°45.0´n 20°41.8´w; map 4). island in the west part of dove bugt. so named by the 1938–39 mørkefjord expedition for the ring-shaped lake in the centre of the island. bonsachs ø has also been used. (ringeøen.) rink mountains 71ø (c. 71°30´n 25°00.0´w). the name is found on maps of lauge koch’s 1926–27 expeditions (koch 1929a), and is de scribed as a wild and jagged range of mountains. it evidently corresponded to the east flank of the stauning alper and the werner bjerge. the name appeared on a geodætisk institut map published in 1931 in a position running diagonally from nordost bugt to antarctic havn, and on a 1932 map from the head of schuchert dal to antarctic havn. a 1952 map placed the name at the werner bjerge. approval of the name by the place name com mittee was suspended pending production of detailed maps, and it was eventually discarded. the name was intended to commemorate heinrich johannes rink [1819–93], a danish greenland explorer and administrator. he was inspector for south greenland from 1855 to 68, and founded the first greenlandic newspaper. risip qaarusaa [hagen] 70ø-217 (70°39.3´n 21°36.5´w). penin sula on the east coast of south liverpool land. one of the names recorded by the 1955 geodætisk institut name registration, it commemorates an occasion when janus sørensen’s assistant ris became so tired on a journey that he had to rest here before he could continue. (risip qârusua, rîsip qârusua.) risip qârusua – see risip qaarusaa. ritomsø 73ø-309 (73°50.1´n 23°10.0´w). elongate lake in central hudson land. named during lauge koch’s 1936–38 expeditions by heinrich bütler after the ritomsee, a large lake in the st. gotthard region of switzerland. rivejernet 71ø-68 (71°35.5´n 25°44.3´w). mountain 2000 m high to the east of borgbjerg gletscher on the north side of nordvest fjord. the name originated from the 1931–34 treårsekspedi tionen, and was approved at the suggestion of ragnar spärck. the name first appeared on the maps of thorson (1934). it was presumably given for its appearance (rivejern = grater). river 1 – river 26 73ø, 74ø (73°55´n to 74°04´n, 21°45´w to 21°55´w). series of rivers in the kap stosch area of home forland, northern hold with hope. the system of numbering rivers was introduced in a report by koch (1931), and as a convenient reference system was subsequently adopted by various other workers (e.g. nielsen 1935; teichert & kummel 1976). some of the rivers also have approved names – foldvik kløft (river 8 or 7), blåelv (river 16), wordie kløft (also river 16), gulelv (river 19), rødelv (river 22) and fosdalen (river 25). see also ekstraelv and river zero (teichert & kummel 1976). river a, d, e, f, g, j 73ø (73°17.5´n 22°34.5´w). reference names used by maync (1942; fig. 18; 1949) and dunbar (1955; fig. 7) for a series of streams east of margrethadal, gauss halvø. river zero 74ø (74°00.8´n 21°54.1´w). river in northern hold with hope, draining west from the western slopes of febold bjerg. this river was originally named ekstraelv by eigil nielsen (1935) during geological work. however, geodætisk institut maps erroneously placed this name against the larger river just to the south, and to avoid confusion the original ekstraelv was renamed river zero. see also discussion in teichert & kummel (1976). the name does not appear on recent official lists of approved names, and is assumed to have unofficial status. rivieradal 80ø-82 (80°03.7´n 21°00.0´w; map 4; fig. 24). e–wtrending valley with several lakes in south kronprins christian land, draining into the south end of hekla sund. so named during lauge koch’s 1952–53 expeditions by erdhardt fränkl, because it is a pleasant valley with relatively rich vegetation. robertselv 74ø (74°10.1´n 20°19.1´w). stream on east clavering ø flowing south into lervig. the name appears on a sketch map in gustav thostrup’s 1921 logbook. robertson ø 73ø-276 (73°04.1´n 23°03.3´w; map 4). island at the mouth of sofia sund, which a.g. nathorst’s 1899 expedition named as robertsons ö after captain tom robertson of the scot 289 tish whaler balaena, which they met several times during the expedition. tom robertson was among the last successful british whalers in east greenland waters, and made regular voyages be tween 1895 and 1907. (robertson island, robertsonøya.) rochusspids 73ø-154 (73°30.7´n 20°27.7´w; map 4). mountain 518 m high in se hold with hope sw of kap broer ruys. named by karl koldewey’s 1869–70 expedition as rochusspitze. the name appeared only on the geological map in koldewey’s narrative, and was not approved until 1939. kommafjeldet has occasionally been used. rock 72ø-6 (72°16.2´n 22°00.7´w; fig. 12). the word rock appears on william scoresby’s 1822 chart against a small island 80 m high off kap young, and was probably intended to indicate a rocky islet rather than a formal name. in the german edition of his narrative (scoresby 1825) it is translated as ‘felsen’. nordenskjöld (1907) combined it mistakenly with another scoresby name to form van dyk rock. cleft island was used by j.m. wordie’s 1926 expedition for the same feature. klippe ø was at one time suggested by the place name committee, but the original ‘rock’ is now the ap proved name. rock lake 77ø (77°35.4´n 20°50.8´w). lake sw of klægbugt, nord marken. named by the 1987 irish expedition to northern east greenland. rold bjerge 72ø-93 (72°44.9´n 23°10.2´w). mountain range on north traill ø, named by ove simonsen during the 1931–1934 tre års ekspeditionen for the danish locality near rebild in jylland. rolige bræ 70ø-8 (70°35.0´n 28°30.0´w; maps 3, 4). glacier on the west side of rødefjord. so named by carl ryder’s 1891–1892 expedition because it seemed to be inactive (rolige = peaceful, quiet). ice bergs at the front of the glacier had not changed their positions between two visits several months apart. rollier bjerge 71ø-167 (71°57.4´n 23°00.1´w). mountain range north of the mouth of ørsted dal. named during lauge koch’s 1936–1938 expeditions by hans stauber after louis rollier [1859– 1931], a swiss palaeotologist who was noted for his studies in the jura and the alps. roma gletscher [ilinnikajia] 70ø-341 (70°03.0´n 22°43.0´w; map 4). glacier on volquaart boon kyst. first explored by leo nardo bonzi’s 1934 expedition, and named ghiacciaio roma after the italian city of rome. the bonzi expedition usage was restricted to the sw branch of the present glacier leading to pyramiden. romer sø 80ø-29 81ø-75 (80°57.0´n 19°27.0´w; maps 1, 4). lake in central kronprins christian land. mapped by lauge koch during flights in 1933 during the 1931–34 treårsekspeditionen, and named probably after the american palaeontologist alfred sherwood romer [1894–1973], noted for his work on permian vertebrates. he was professor at harvard university from 1934 to 1963, and director of the museum for comparative zoology from 1946 to 1961. (romer lake.) romeydalen 74ø-341 (74°45.7´n 20°03.7´w). valley on se kuhn ø. named during lauge koch’s 1936–38 expeditions by wolf maync and andreas vischer. rommelshausener spids 71ø (71°50.8´n 25°16.8´w; map 5). moun tain on the sw side of roslin gletscher. climbed by karl herlig koffer’s 1966 scoresby land expedition on 21 august, and named after rommelshausen north of stuttgart, the home town of günter schnaidt, one of the three climbers. (rommelshausenerspids.) ronicol 71ø (71°38.4´n 25°18.4´w; map 5). high pass between ox ford gletscher and the head of jupiter gletscher, south stau ning alper. crossed by the 1975 scottish expedition, and named apparently for a brand of frostbite tablets. roon bugt 76ø-4 (76°18.0´n 20°00.0´w; map 4). bay on the east side of hestefoden, between teufelkap to the north and ad. s. jensen land to the south. named by karl koldewey’s 1869–70 expedition as roonbai, after albrecht von roon [1803–1879], prus sian field marshal and minister of war, who was present at bremerhaven when the expedition sailed in 1879. koldewey’s usage was much broader than the present, and included much of what is now the fig. 73. the 2783 m high mountain rigny bjerg inland from the blosseville kyst. it was first seen during the norse voyages from iceland to southwest greenland, and features in the icelandic sagas as bláserk. it was named mt. rigny (now rigny bjerg) by jules de blosseville in 1833. aerial photograph, © geodætisk institut. rigny bjerg 290 southern part of dove bugt. (roon-bai, roon bay, roons bugt.) roscoe bjerge 70ø-175 (70°40.7´n 22°01.3´w). mountain chain in south liverpool land, thickly crested and serrated with pinnacles, whose north limit is taken as a line through sødal and aage nielsen gletscher. the mountains were named by william scoresby jr. in 1822 as the roscoe mountains in compliment to william roscoe [1783–1831], an historian who became mp for liverpool in 1806. the name did not appear on maps until 1934, when its usage was reinstated at the suggestion of brian roberts. (roscoe berge.) roseneath – see mønstedhus and ottostrand. roseneathbugt 75ø-3 (75°42.9´n 19°31.0´w; map 4). pronounced bay on the north part of the east coast of hochstetter forland, south of haystack. named by douglas clavering in 1823 in the form roseneath inlet after the locality opposite ardencaple castle, dumbarton, scotland (rosneath on modern maps). it is the site of a castle, now demolished. clavering apparently could not see the flat area of hochstetter forland, and his name was probably originally applied to the area around agnete sø which looked like a fjord. karl koldewey’s 1869–70 expedition first applied the name to the present locality south of haystack. (roseneath bay.) rosenheimer spids 71ø (71°53.9´n 25°27.2´w). mountain 1950 m high on the ridge between duart gletscher and the upper basin of spærregletscher. climbed by karl herligkoffer’s 1966 expedition, and named after the small town of rosenheim at the foot of the bavarian alps. it has also been called piz coaz. rosenjoch 75ø (c. 75°19´n 17°50´w). highest point in the camp area of the 1943–44 operation bassgeiger at kap sussi, shannon. the name is reported by olsen (1965). rosenvinge 70ø (70°29.1´n 21°57.9´w). name occasionally used in error for the town of scoresbysund in reports by french scientists of j.-b. charcot’s expeditions (e.g. faure 1933). scoresbysund lies on the north side of rosenvinge bugt. rosenvinge bugt 70ø-304 (70°27.6´n 22°05.0´w). large bay in south liverpool land. named by g.c. amdrup’s 1898–1900 expedition, after janus andreas kolderup-rosenvinge [1858–1939], a danish botanist and professor at the university of copenhagen. (rosenvinge bay, bai de rosenvinge, rosenving bay, rosenvinge baie.) rosinante 73ø-424 (73°21.0´n 25°07.9´w; fig. 74). mountain 758 m high on nw ymer ø. named during lauge koch’s 1947–49 expeditions by silvio eha, possibly for a supposed resemblance to don quixote’s horse. it has also been called little chocolate mountain. rosinante pas 73ø-423 (73°22.6´n 25°06.8´w). minor pass on nw ymer ø, ne of rosinante. named during lauge koch’s 1947–49 expeditions by silvio eha. rosio 77ø-13 (77°17.0´n 18°21.8´w; map 4). small island on the south side of skærfjorden. named by the duke of orléans in 1905 as ilot del rosio. the origin of the name is unknown. the name was mis placed by the 1906–08 danmark-ekspeditionen, the original ro sio being the present rekvedøen. the wrong position had been extensively used in reports of the 1906–08 danmark-ekspedi tionen before the error was discovered, and it was considered that correction would only have caused confusion. (rosio ø, ile del rosio.) roslin borg 71ø-312 (71°54.0´n 24°17.5´w; map 5). mountain 2560 m high at the head of roslin gletscher, south stauning alper. it was first climbed by malcolm slesser’s 1958 expedition, and named after roslin castle, near edinburgh, part of which dates from the 14th century. (roslinborg.) roslin gletscher 71ø-313 (71°48.0´n 24°48.2´w; maps 4, 5). glacier more than 20 km long in the south stauning alper, flowing se to schuchert dal. it was first traversed by malcolm slesser’s 1958 expedition and named roslin glacier, officially approved in danicised form in 1959. due to inaccurate topographic maps and some confusion it was some years before it was realised that this same glacier had been officially named ivar baardsøn gletscher in 1939. the latter name had rarely been used on maps, and was discarded in 1971 in favour of roslin gletscher. rosmule 76ø-127 (76°39.4´n 24°22.5´w; map 4). peninsula on the south side of borgjøkel, dronning louise land, named by j.p. koch’s 1912–13 expedition. possibly given for the shape (rosmule = horse muzzle), or an incident with the horses used on the expedition. (rosmulen, rossmule.) rossily bjerg 71ø-22 (71°55.5´n 22°44.7´w). mountain 770 m high in east scoresby land, originally named cape rossily by william scoresby in compliment to a french philosopher, probably fran cois etienne rosily-mesros. scoresby used the form cape rossilly in the appendix of his english (1823) narrative, and in the german (1825) edition spelt it variously ‘rosilly’, ‘rossilly’ and ‘rossily’. rosily and de rossel had written a report on scoresby’s ‘account of the arctic regions’ for the french government. (cape rosilly.) rostrum avis 76ø (76°57.8´n 20°33.1´w). name used for the feature fuglenæbsfjeldet on the christmas card sent to peter freuchen at pustervig in 1907. it is reproduced in koch (1916 p. 398). (fugle næb = bird’s beak = rostrum avis). rothé island 70ø (70°52.3´n 21°40.0´w). name proposed for the present janus ø off the coast of south liverpool land by helge g. backlund during the 1931–34 treårsekspeditionen. jean rothé was a french geophysicist, one of the party manning the french international polar year station at scoresbysund in 1932–33, who joined backlund’s party for a few days in july 1933. the name is found only in kranck (1935). (rhoté i.) round pond 72ø (72°14.4´n 23°53.9´w). name used by the 1968–74 dundee university expeditions for a small pool near langdyssen at the ne end of mestersvig airfield, possibly identical with gåsesø. rovmågesø 70ø-310 (70°29.4´n 21°53.4´w). lake east of scores bysund town in south liverpool land. named during the 1924–25 colonisation expedition for the numerous arctic gulls (pedersen 1926). (rovmaagesø, arctic gull lake.) royal peak 72ø (72°04.2´n 24°44.0´w; map 5). mountain 2500 m high between the head of bersærkerbræ and schuchert gletscher, stauning alper. it was first climbed by the 1961 bangor expedition. the second ascent has been stated to be by the 1963 imperial college expedition (bennet 1972), which called it westminster. however, some climbers consider westminster to be a subsidiary summit a short distance east of royal peak. royston nunatakker 71ø-66 (71°22.5´n 29°42.0´w; maps 3, 4). nunatak group south of daugaard-jensen gletscher. mapped by lauge koch during flights in 1932 on the 1931–34 treårseks peditionen, and apparently named after the small town of royston north of london where arthur hinks had a summer cottage (see also hinks land). rubjerg knude 72ø-81 (72°42.1´n 23°33.1´w; map 4). mountain on north traill ø. named by ove simonsen during the 1931–34 treårsekspeditionen after the danish locality of the same name on the coast sw of hirtshals, jylland. rud.-johansen valley 72ø (72°56.2´n 25°21.9´w). valley in se suess land, the present ørkendal. the name was used by eugéne weg mann during the 1931–34 treårsekspeditionen (wegmann 1935), and was in general use by expedition members in 1931–1932. it was said to be named after the master-tailor rud.-johansen, who had made donations to the expedition. rudbeck bjerg 73ø-275 (73°01.8´n 23°17.9´w; map 4). mountain 1322 m high on north geographical society ø. named by a.g. nathorst’s 1899 expedition as rudbecks berg, possibly after olaf (olaus) rudbeck [1660–1740], a noted swedish scientist, whose best-known work was his five-volume history of sweden published in 1679. (rudbeck mountain, rudbeckfjellet, rüdbeckberg, rudbeck bjærge, rudbecktinden.) rudi bugt 74ø-119 (74°23.4´n 21°45.6´w; map 4). small fjord on the nw side of clavering ø, named by lauge koch’s 1929–30 expeditions in the form rudi bay for the norwegian hunter henry rudi who hunted from moskusheimen (also known as revet) at the 291 andrée land noa sø rosinante chokoladebjerg grejsdalen head of the bay for many years. henry rudi was one of the most successful of norwegian hunters, and was known as ‘isbjørn kongen’ (= the polar bear king). in the course of his hunting career in east greenland and svalbard he shot 713 polar bears, including 113 in a single year in svalbard. he is also reputed to have shot 70 falcons in east greenland in the autumn of 1941. (rudifjorden.) ruin ø 71ø (71°15.7´n 24°55.8´w). name used by glob (1946) for the islands se of sydkap also known as immikkeertivaqqat, where a large inuit settlement with 10 house ruins was found by helge larsen in 1937. ruinelven – see gravelven. ruinerne 71ø-269 (71°57.5´n 23°58.9´w; map 5). mountain 1314 m high in the werner bjerge north of the head of sirius gletscher, named by peter bearth and eduard wenk during lauge koch’s 1953–54 expeditions (ruinerne = the ruins). it was climbed by bearth in 1953. rumpen 73ø-429 (73°16.7´n 24°48.1´w; map 4). isolated hill on the south side of inner dusén fjord, ymer ø. the name was modified by the place name committee from a suggestion by silvio eha. eha (1953) used a greenlandic name nulog on his cross-sections. both names record the shape (rumpen = the rump, behind). rund top 72ø (72°51.0´n 22°27.8´w). mountain 726 m high on east geographical society ø behind kap mackenzie, the present leitch bjerg. the name was used on one of the map’s of carl ryders’s 1891–92 expedition, probably in a purely descriptive sense (rund top = rounded summit). rundefjeld 70ø-10 (70°31.7´n 28°36.3´w; map 4). rounded icecapped summit 1512 m high on the landmass between rolige bræ and vestfjord. named by carl ryder’s 1891–92 expedition as runde fjæld. rundetårn 74ø-114 (74°10.9´n 20°30.2´w). mountain 830 m high on east clavering ø. named by lauge koch’s 1929–30 expeditions in the form mt. rundetaarn after the church tower and observatory of the same name in copenhagen. (rundetaarn berg, runde taarn, runde taarn bjærg.) rundgletscher 72ø-348 (72°14.3´n 22°40.1´w). glacier on se traill ø. named by h.p. heres during lauge koch’s 1956–58 expeditions for its shape. rundholmen 73ø (73°03.9´n 22°33.2´w). small island in the broch øer group, so named on the 1932a nsiu map for its round shape. rundsø 72ø-124 (72°52.5´n 25°08.3´w). circular lake on nw ella ø, so named during the 1931–34 treårsekspeditionen by the ella ø wintering party. (round lake, rundsee.) rundvika 72ø (72°55.6´n 22°02.7´w). bay on east geographical society ø, nw of kap mackenzie. used on the nsiu maps of lacmann (1937) and named for its round shape. runetind 73ø (73°06.1´n 23°40.3´w). small isolated peak on a mountain ridge in se ymer ø. so named on the 1932a nsiu map, probably for markings resembling runes. rungsted elv 72ø-197 (72°12´n 24°00´w; map 5). river draining the flanks of korsbjerg and domkirken, north scoresby land, named after the village of rungsted north of copenhagen, den mark. rungstedbjerg 74ø-406 (74°02.4´n 22°38.4´w). mountain in the nørlund alper, north of rungstedgletscher, north hudson land. named after the village of rungsted near copenhagen. this name appears to have been suggested by the place name committee as a substitute for vermessungsbjerg, used by helge backlund for the same mountain. rungstedgletscher 74ø-134 (74°01.7´n 22°40.6´w; map 2). gla cier in the nørlund alper, north hudson land. named by lauge koch’s 1929–30 expeditions in the form rungsted glacier, for the danish locality (see rungsted elv). (rundstedbreen.) rustplateau 74ø-173 (74°11.7´n 21°19.6´w). plateau on sw clavering ø, culminating to the north in vestmar bjerg. the name was first used in reports of the 1931–34 treårsekspeditionen in the form rust plateau (malmquist 1932), and refers to the vivid red and yellow rusty colouration due to weathering of disseminated pyrite. associated veins contain small amounts of galena, sphalerite, chalcopyrite and pyrite. ruth ø 72ø-46 (72°59.6´n 24°53.1´w; map 4). island east of suess land. named by a.g. nathorst’s 1899 expedition as ruths ö (fig. 8), after his daughter who had her 16th birthday the day kong oscar fjord was discovered. ruth gabriella nathorst [b. 1883] was a missionary in china from 1918 to 1944. see also maria ø and ella ø. (ruth island, ruthöya.) rutherford bjerg 77ø-126 (77°05.0´n 24°35.2´w; map 4). highest mountain in the mountain range south of krebs bjerg, dronning louise land. one of the names given by the 1952–54 british north greenland expedition for notable scientists, it commemorates the fig. 74. view northwards across western ymer ø, with noa sø, rosinante and chokoladebjerg in the foreground, and andrée land and grejsdalen in the background. the john haller photograph collection, geus archive. 292 british physicist lord rutherford [1871–1937]. he laid the ground work for the development of nuclear physics, and had an influence on scientific thought comparable to faraday and newton. ruthner 74ø-27 (74°51.8´n 20°00.0´w; map 4). mountain 1060 m high on east kuhn ø. so named by karl koldewey’s 1869–70 expedition, probably after anton von ruthner [1817–97], a noted aus trian mountaineer (j. løve, personal communication 2010). (ruth ner berg.) ruthven spids 72ø-359 (72°02.5´n 25°09.9´w; map 5). mountain 2400 m high on the ne side of sefström gletscher, north stauning alper. first climbed by malcolm slesser’s 1958 expedition, and named after ruthven barracks, near kingussie, scotland, used by english troops in the 17th century and now preserved as an ancient monument. (ruthven.) ryder elv 70ø-119 71ø-124 (71°00.0´n 22°29.1´w). river occupying the valley between jameson land and liverpool land. it was originally seen by carl hartvig ryder [1858–1923] from neill klin ter in 1891–92, a discovery that demonstrated that hurry inlet was a fjord and not a sound. the name originated from a.g. nathorst who named ryders älf when he visited and mapped the end of the fjord in 1899. (ryder river, ryder elv.) ryders dale 70ø, 71ø (71°00.0´n 22°29.1´w). name employed by nordenskjöld (1907) for the present klitdal between liverpool land and jameson land in which ryder elv flows. named after carl hartvig ryder. see also ryder elv. (ryder’s valley, vallé ryder.) ryders depot 70ø (c. 70°27´n 22°37´w). house built at kap stewart by carl hartvig ryder in 1892, where a depot was left for possible emergency use by subsequent visitors. the name first appears on the maps of the g.c. amdrup’s 1898–1900 expedition. in 1924, the house was reported damaged by ejnar mikkelsen and the depot pillaged. he rebuilt the house and added two additional houses for the settlement of kap stewart or ittorisseq. slogans painted on the hut recorded visits by the scotia of dundee in 1905, the quest, and the bonø in 1924. the house was subsequently demolished. ryledammen 74ø (74°28.9´n 20°31.5´w). small pond in the eastern part of rylekærene, east of zackenberg forskningsstation. the name is used as a reference locality in ornithological reports of visiting scientists. (ryledammene.) rylekær 76ø-240 (76°49.6´n 19°05.8´w). boggy area on winge kyst where sandpipers (= ryle) were commonly observed. so named in the ornithology reports of the 1906–08 danmark-ekspe di tionen. rylekærene 74ø (74°29.1´n 20°31.7´w). area ne of zackenberg forskningsstation, where sandpipers (= ryle) commonly nest. the name is used as a reference locality by visiting scientists. rypedalene 70ø (70°29.0´n 26°17.3´w). small valleys or ravines in the rypefjeldene on danmark ø. the reference is only found in the report by hartz (1895) on work during the wintering of carl ryder’s 1891–92 expedition in hekla havn. rypefjeldene 70ø-61 (70°29.3´n 26°18.0´w). low hills on the north side of elvdalen on danmark ø. named by carl ryder’s 1891–92 expedition as rypefjældene. they were notably free of snow in the winter and the haunt of ptarmigan (= rype). (rype fjældene.) rypefjeldet 76ø-55 (76°56.2´n 20°22.1´w). hill on the west side of mørkefjord station, south of sælsøen. so named by the 1906–08 danmark-ekspeditionen because ptarmigan were shot here in 1906 (fig. 75; thostrup 2007). (ptarmigan hill, rype mt., rjúpna fell.) rypefjord [aqissip kangertiva] 70ø-3 71ø-38 (71°00.0´n 27°40.0´w; maps 3, 4). fjord between sw renland and c. hof mann halvø. named by carl ryder’s 1891–92 expedition, which discovered the fjord during a sledge journey in april 1892. ptarmigan (lagopus mutus) are common in the region (fig. 75). the greenlandic name has also been recorded as aqissit kanger suat. (rype fjord.) rypegletscher 73ø-546 (73°00.5´n 28°07.6´w). glacier in north goodenough land, flowing nw to join nordenskiöld gletscher. named by j.m. wordie’s 1929 expedition as ptarmigan glacier, because occasional ptarmigan were seen here. the glacier was divided into three parts, upper, middle and lower. two of these have official names, øvre ptarmigangletscher and nedre ptarmi gan gletscher. (ptarmigan glaciers, ptarmigangletscher.) rypenæs 71ø-325 (70°59.4´n 27°42.4´w). peninsula on the west side of rypefjord. the name was approved in 1961 at the suggestion of ulrik røen. recent 1:100 000 scale topographic maps show the location to be just south of latitude 71°n. rypesletta 74ø (74°29.6´n 19°00.0´w). the name has been used by norwegian hunters for the plain west of the hunting station at kap wynn, where there always seemed to be ptarmigan. rypesvinget 74ø (74°28.5´n 20°34.1´w). pronounced bend in the river north of zackenberg forskningsstation. the name is used by visiting scientists. rypesø 72ø-199 (72°13.6´n 23°55.4´w; map 5). small lake west of noret, north scoresby land. named by prospecting teams associated with lauge koch’s 1948–49 expeditions. rytterknægten 72ø-109 (72°47.9´n 25°15.7´w). cape in ne lyell land on the sw side of narhvalsund. the name was given by the place name committee in 1935, as a replacement for an unsuitable proposed name, probably after the highest point on the danish island of bornholm. rævebræ 70ø-430 (70°09.5´n 26°54.5´w). small glacier on the south side of gåsefjord. so named during the 1967–72 ggu scores by sund expeditions by e.a. hailwood because of the many foxes (= ræve) seen in the vicinity (fig. 76). rævedal 73ø-358 (73°50.7´n 24°38.1´w). valley in east strindberg land, draining east to waltershausen gletscher. named during lauge koch’s 1948–49 expeditions by hans r. katz, for the arctic fox (fig. 76). ræveelv 70ø-115 (70°55.4´n 22°51.0´w). river in east jameson land west of the head of hurry inlet, draining into ugleelv. named fig. 75. ptarmigan (rype) in summer plumage, a common bird throughout east greenland. photo: jakob lautrup. 293 by alfred rosenkrantz during lauge koch’s 1926–27 expeditions as fox river. ræveelv 74ø (74°29.6´n 20°33.0´w). minor stream near to zacken berg forskningsstation. the name is used as a reference locality by visiting scientists. rævehalen 70ø-420 (70°40.9´n 29°18.8´w). nunatak on the north side of the upper part of roligebræ. so named by laurent jemelin during the 1967–72 ggu scoresby sund expeditions because of the shape on the map, somewhat like a fox tail. rævehøjene 74ø (74°29.6´n 20° 35.6´w). hill in the vicinity of zac ken berg forskningsstation. the name is used as a reference locality by visiting scientists. rævekløft 70ø-282 (70°27.2´n 22°37.1´w). gulley near kap ste wart in se jameson land. named during the 1924–25 expedition that founded scoresbysund for the numerous foxes (pedersen 1926). a total of 25 were observed in the vicinity of ryders depot in october 1924. rævekløft 74ø-220 (74°01.1´n 21°31.8´w). minor ravine in nw hold with hope, between rivers 9 and 10, on the north slope of frebold bjerg. so named during the 1931–34 treårsekspeditionen by eigil nielsen. rævestenen 74ø (74°28.0´n 20°30.9´w). rock close to kærelv, east of zackenberg forskningsstation. the name is used as a reference locality by visiting scientists. röbeckstua 73ø (73°53.2´n 20°18.2´w). norwegian hunting hut at kap james, ne home forland, built by the møre expedition in august 1930. the name is found in this form on the 1932a nsiu map, and as rœbeck-stua on other nsiu maps (1932c). it was named after the brothers peder and knut røbek, both hunters with the expedition. knut røbek was drowned when he fell through the ice off the south coast of clavering ø in december 1931, and his grave is on the hill behind herschellhus. the hut has also been known as kap james hytten. (røbekstua.) røbekfjellet 74ø (74°09.0´n 21°03.4´w). mountain c. 1080 m high on south clavering ø. named on nsiu maps of lacmann (1937), after peder røbek [b. 1897], a norwegian hunter who wintered in east greenland in 1927–29 and 1930–31. røde bakker 81ø (81°18.8´n 13°50.2´w). hills in nw kilen, kron prins christian land, where a wine-red marker layer of upper cre ta ceous age crops out. the name is found on a coloured geological map of kilen printed in 1991 (pedersen 1991). røde elvdal 70ø-2 (70°58.7´n 28°05.0´w). valley running into harefjord. so named by carl ryder’s 1891–92 expedition because red (= rød) conglomerates were found here during a sledge journey in april 1892. røde fjæld 76ø (76°10.2´n 18°39.8´w). name used by friis (1909) in his popular account of the 1906–08 danmark-ekspeditionen for a red-coloured mountain in one of the ravines crossing store kolde wey, probably that north of trækpasset. this may be identical with the mountain that danish hunters and personnel at danmarkshavn weather station know by the same name. røde hytte [aappalaartukajik] 70ø (70°33.7´n 23°44.3´w). hun ting hut on the coast of sw jameson land. the name has been used in archeological reports. the hut was originally painted red. røde mur 71ø-258 (71°59.0´n 24°08.3´w; map 5). mountain ridge in the werner bjerge on the ne side of langefirn. named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk for its colour (= red wall). quartz-fluorite veins in the east part of the wall contain abundant pyrite, the rusty weathering of which is responsible for the colour. røde roseelv 73ø-291 (73°56.7´n 22°05.0´w). river in east hud son land draining east into loch fyne. the name was used by gunnar säve-söderbergh during the 1931–34 treårs ekspedi tionen, originally in the form red rose valley. røde støvhorn 72ø-422 (72°52.9´n 27°04.3´w). mountain on the north side of dickson fjord. named during the 1931–34 treårs eks pe ditionen by eugéne wegmann originally as red staubhorn, after a professor staub of zürich, an ironic tribute to a colleague who as a consequence of excessive drinking often had a red nose. the mountain was climbed by wegmann’s party on 4 august 1932. rødebjerg 72ø-214 (72°08.4´n 24°01.3´w; map 5). mountain ridge sw of mesters vig. named by prospecting teams associated with lauge koch’s 1948–49 expeditions for the colour. rødebjerg 73ø-529 (73°03.7´n 24°18.0´w; map 4). mountain 1683 m high on south ymer ø. named as röda berget by a.g. nathorst’s 1899 expedition because it was made up of red devonian sandstone. (red mountain, røda mountain, røda berget.) rødebjerghytten 73ø (73°02.8´n 24°04.7´w). norwegian hunting hut built for arktisk næringsdrift in october 1929 on the north side of sofia sund, about 7 km east of rødebjerg. it has also been known as arentzhytta and snehytten. rødedal 72ø-376 (72°00.9´n 23°42.0´w). valley on the sw side of oksehorn, draining into kolledalen, north scoresby land. the name was used by hans kapp during lauge koch’s 1957–58 expefig. 76. arctic fox (ræv) in summer. foxes feed on lemmings and young birds. photo: jakob lautrup. 294 ditions. officially it is considered to be identical with bearth’s lille oksedal, but kapp evidently considered the latter to be a minor side valley to rødedal. rødedal 73ø-356 (73°52.4´n 24°55.4´w). valley in central strind berg land, draining south to join brogetdal. named during lauge koch’s 1948–49 expeditions by hans r. katz. rødedal 73ø-50c (73°58.1´n 21°23.4´w). minor valley on the north slope of stensiö plateau, nw hold with hope, draining into blåelv. named by eigil nielsen during the 1931–34 treårsekspeditionen for the colour of the rocks. rødedalen 73ø (73°18.7´n 22°26.8´w). valley between knolden and saxo bjerg, in the giesecke bjerge, draining ne into margrethe dal. named during lauge koch’s 1936–38 expeditions by maync (1942, 1949). the valley is carved into red devonian rocks. rødefjeld 73ø-414 (74°02.6´n 28°27.3´w). red coloured nunatak in arnold escher land, named by hans r. katz during lauge koch’s 1951 expedition. (rödefjeld.) rødefjord 70ø-24 (70°45.0´n 27°50.0´w; maps 3, 4). fjord west of milne land named røde fjord by carl ryder’s 1891–92 expedition. the west side of the fjord is largely formed by conspicuous cliffs of red conglomerate. (röde fjord, røde ø fjord.) rødelv 70ø-120 (70°56.0´n 22°33.8´w). river at the head of hurry inlet draining from dusén bjerg into ryder elv. named by alfred rosenkrantz during lauge koch’s 1926–27 expeditions as red river, for the colour. rødelv 73ø-49 (73°55.1´n 20°56.3´w). river in north hold with hope draining north into gael hamke bugt. named by lauge koch’s 1929–30 expeditions as red river, for the colouration due to the triassic rocks. it corresponds to river 22 of koch (1931). a norwegian hut about 1 km se of the river sometimes known as rødelv, is more usually known as knophstua. (raudelva.) rødelv 75ø-56 (75°12.2´n 20°00.0´w). river in south hochstetter forland, named during the 1931–34 treårsekspeditionen by hans frebold in the form roten bach (= red river), for the colour. rødepynt 70ø-4 (70°51.3´n 27°53.2´w). cape at the junction of harefjord and rødefjord, named røde pynt by carl ryder’s 1891– 92 expedition because of the cliffs of intense red conglomerate which begin here. (røde pynt, røde punkt.) rødevæg 76ø-219 (76°52.5´n 21°03.3´w). mountain wall on the north side of hellefjord. named by the 1938–39 mørkefjord expedition, for the red granitic rocks. rødeø 70ø-21 (70°27.7´n 28°05.0´w; map 4). island at the south end of rødefjord. named røde ø by carl ryder’s 1891–92 expedition because it was composed of red conglomerate. (röde o, røde island, red island.) rødeø 76ø-27 (76°43.1´n 20°55.4´w; map 4). island on the west side of dove bugt, named by the 1906–08 danmark-ekspedi tionen as røde ø, because it was entirely composed of conspicuous red granite. (röde ö, røde island.) rødhorn 72ø-453 (72°53.1´n 26°52.4´w). mountain on the north side of dickson fjord, so named during the 1931–34 treårsekspe di tionen by eugéne wegmann who climbed it on 6 august 1933. rødkam 73ø-376 (73°46.7´n 26°06.8´w; map 4). mountain in andrée land south of eremitdal, named by erdhardt fränkl during lauge koch’s 1948–50 expeditions for the colour of the rocks. rødkilefjeld 70ø-441 (70°28.9´n 29°02.6´w). mountain 1660 m high between rolige bræ and vestfjord. named by laurent jemelin during the 1967–72 ggu scoresby sund expeditions for the wedge of red rocks at the foot of the mountain. rødryggen 74ø-346 (74°32.1´n 19°49.7´w). low ridge in wolla ston forland, formed by red coloured sedimentary rocks. named during the 1936–38 two-year expedition by wolf maync. (roter rücken.) rødstak 71ø-439 (71°43.9´n 24°13.4´w). mountain 1276 m high with a summit of red rocks in the gurreholm bjerge east of schuchert dal. the name was first used by rudolf trümpy, and was authorised at the suggestion of ggu in 1974. rødsten 73ø-102 (73°21.0´n 24°03.0´w). mountain on the north side of gunnar andersson land, ymer ø. named during the 1931–34 treårsekspeditionen by th. johansen for the red colour. rødtop 73ø-94 (73°48.2´n 23°57.0´w). mountain in sw hudson land north of the mouth of moskusoksefjord, named during the 1931–34 treårsekspeditionen by th. johansen for the redcoloured summit. rødtophytten 73ø (73°48.4´n 24°02.2´w). norwegian hunting hut built by arktisk næringsdrift in 1938 on the east side of walter hausen gletscher at the foot of rødtop. it is also known as brehytten and solstrand. rødøen 74ø-88 (74°07.3´n 22°52.7´w). semi-nunatak 913 m high on the south side of wordie gletscher, named by j.m. wordie’s 1926 expedition as red island for its colour. it was thought to be an island-like nunatak surrounded by ice, but later found to be bounded by a valley on its sw side. røgelen 73ø-198 (73°46.2´n 21°39.6´w). mountain 552 m high on the east side of loch fyne. named on an nsiu map (1932a) in the form rugelen, possibly because of the ominous manner in which clouds developed. the approved danish form suggests ‘smoking’. röhlingfjellet – see mont röhling. röhss fjord 72ø-403 (72°44.0´n 26°37.0´w; map 4; see also fig. 52). e–w-trending fjord which divides gletscherland almost into two parts. named röhss' fjord by a.g. nathorst’s 1899 expedition after johan anders august röhss [1836–1904], who contributed 5000 swedish kronor to the expedition finances. röhss was a successful swedish merchant and an important figure in the social life of gothenburg. (röhs fjord, røhss fjord, røhs fjorden, røss fjord.) röhss fjord hytten 72ø (72°42.4´n 26°47.6´w; see also fig. 82). nor wegian hunting hut built for arktisk nærings drift in july 1934 on the east side of strømnæs, röhss fjord. it was originally known as festningen, and is often called strømnæshytten. röhssfjordbotnen 72ø (72°40.3´n 27°11.8´w). this name has been used in nsiu botanical reports as a reference locality for the innermost section of röhss fjord (vaage 1932). röhssfjordsundet 72ø (72°42.6´n 26°50.0´w). name used as a reference locality in nsiu botanical reports (vaage 1932) for the narrow part of röhss fjord at strømnæs known as sarpaq. røiskattlia 73ø (73°38.9´n 23°10.5´w). norwegian hunting hut built in august 1932 for arktisk næringsdrift on the north side of moskusoksefjord, commonly known as petrahytten. the area was plagued by ermine (= røyskatt), which often broke into the hunters’ fox traps. the hut has also been known as første hytten. røiskattlia 73ø (73°11.1´n 25°58.4´w). norwegian hunting hut built by arktisk næringsdrift in july 1947 on the north side of suess land, and commonly known as polarheimen. the name records the presence of ermine (= røyskatt). røievatnet 74ø (74°00.0´n 22°04.0´w). small lake on the west side of outermost loch fyne. the name was used by nsiu in the 1930s for a lake where they fished for salmon (arctic char = røie). it has been used as a reference locality in botanical and zoological reports. (røyevand, röyevatnet, røyevatnet, røjevand, røjevandet.) rømer fjord 69ø-20 (69°44.0´n 23°36.0´w). fjord sw of turner ø on the north blosseville kyst. named by g.c. amdrup’s 1898– 1900 expedition as rømers fjord, after ole rømer [1644–1710], a danish physicist and astronomer noted especially for the first measurement of the speed of light. røseløbet 76ø-80 (76°40.5´n 18°43.7´w). narrow sound between the two islands of lille koldewey. so named by the 1906–08 dan mark-ekspeditionen because it is only a sound at spring tides (thostrup 2007). (röselöbet, röse löbet.) røsnes 74ø (74°42.4´n 20°03.8´w). name occasionally used by nor wegian hunters for kap hamburg, southern kuhn ø. the name has also been used for the norwegian hunting hut 3 km west of the cape 295 commonly known as furnes. røstholmane 72ø (72°42.2´n 21°50.6´w). small skerries off the coast of se geographical society ø. used only on nsiu maps (lacmann 1937), the name was given for the island røst in the lofoten region of norway. (röstholmane.) røvballehytten 72ø (72°59.0´n 24°33.4´w). name often used for the norwegian hunting hut built for arktisk næringsdrift in sep tem ber 1930 in inner sofia sund, which is also known as svedenborg, bakkehytta, joplassen and valborghytta. the name stems from the position of the hut on a steep slope, which was difficult to reach with a heavy sledge and dogs. røvdalen – see raudalshytta. røverreden 70ø-79 (70°13.6´n 25°01.2´w; map 4). mountain on the north side of bredegletscher, a fantasy name (= den of thieves) given by laurits bruhn during the 1931–34 treårsekspeditionen. råumøyane 72ø (72°44.2´n 22°50.7´w). two islands in vega sund, part of the scott keltie øer group. so named on nsiu maps of lac mann (1937), after knut råum [b. 1909], a norwegian hunter who wintered in east greenland in 1933–35 and 1936–37. s s. paolo 72ø (72°04.4´n 25°07.4´w). one of the pinnacles of satans galleri in the stauning alper, nne of korsspids. climbed on 29 july 1984 by sandro pucci’s expedition, and named after two of the climbers (paolo piconi and paolo d'ugo). s. thomsen pynt 76ø-91 (76°42.0´n 18°29.8´w). south cape of joh. g. guildal ø. named by the 1906–08 danmark-ekspedi tionen as s. thomsens pynt, possibly after sigismund gotthelf thom sen [1831–1903] who was thomas thomsen’s father. see also thomas thomsen næs. saakattaakajik 71ø-224 (71°17.8´n 24°54.4´w). point east of syd kap on the west side of nordøstbugt. recorded by the 1955 geo dætisk institut name registration, the name means ‘the rather flat’. (sâkátâkajik.) sabine ø 74ø-54 (74°35´n 18°56´w; maps 2, 4). island ne of wol la ston forland, one of the pendulum øer. named by karl kolde wey’s 1869–70 expedition as sabine insel (fig. 6), after edward sabine [1788–1883], british general and physicist who carried out pendulum experiments on the island in 1823. it has also been called inner pendulum island. (sabine island, sabineön.) saddelfjeld 81ø (81°20.2´n 14°05.1´w). hill 419 m high in nw kilen, kronprins christian land, where saddle-shaped folded strata dominate. the name is found on a coloured geological map of kilen printed in 1991 (pedersen 1991). saddle lake 77ø (77°34.1´n 20°48.8´w). lake on a col south of klægbugt, nordmarken. named by the 1987 irish expedition to northern east greenland. sadelbjerg 74ø-60 (74°23.7´n 19°36.8´w). mountain 1181 m high with twin summits in wollaston forland. named by karl kolde wey’s 1869–70 expedition as sattelberg, for its saddle-like shape. the first ascent was made by julius payer and ralph copeland in september 1869. (sattelberges, salfjellet, mt saddelbjærg, saddle moun tain, granitsattelberg.) sadelen 73ø-379 (73°43.9´n 25°20.7´w). col or pass between mo ræne dal and geologfjord. named during lauge koch’s 1948–50 expeditions by erdhart fränkl. it is said to be the easiest place to reach for a good view of inner geologfjord (sadelen = the saddle). (sadlen.) sadlen 76ø (c. 76°56´n 21°03´w). mountain on the south side of mørkefjord, west of danmarksmonumentet. the name is found on charles poulsen’s (1991) map (j. løve, personal communication 2009). säve söderbergh bjerg 73ø-444 (73°26.9´n 22°36.7´w). mountain on eastern gauss halvø, south of agassiz bjerg. the name is attributed to heinrich bütler, and arose during his work with lauge koch in the 1950s. it commemorates gunnar säve-söderbergh [1910–1948], a swedish palaeontologist who participated in lauge koch’s east greenland expeditions from 1932 to 1936, and made studies especially of devonian and triassic rocks. sâkátâkajik – see saakattaakajik. salèvebjerg 73ø-317 (73°40.5´n 22°30.6´w). mountain in se hud son land, so named during lauge koch’s 1936–38 expeditions by heinrich bütler after mont salève, a mountain se of geneva, switzerland. (salèveberg.) salix dal 70ø-405 (70°41.6´n 23°19.9´w). minor valley in sw jameson land draining into sjællandselv. named during the 1967– 72 ggu scoresby sund expeditions by tove birkelund, for the unusually rich willow vegetation. sandbach halvø 70ø-229 (70°44.2´n 21°38.6´w). prominent penin sula between vejle fjord and kolding fjord in liverpool land. named sandbach island by william scoresby jr. in 1822 after a much respected friend. the ‘island’ was later found to be a peninsula (= halvø). (sandbach ö.) sanddal 78ø-47 (78°04.7´n 21°35.4´w; map 4). valley in hertugen af orléans land. so named during lauge koch’s 1956–58 expeditions by john haller, because the valley was full of sandy alluvium. sanddalen 73ø (73°00.3´n 23°53.7´w). valley on the north side of geographical society ø, so named by nsiu in 1930 because of the deposits of sand at its mouth. norwegian and danish botanists have used the name as a reference locality. (sanddal.) sanderling river 72ø (72°30.8´n 23°58.9´w). name used by uni versity of dundee expeditions between 1968 and 1974 for a minor stream west of karupelv draining into holm bugt, sw traill ø. it was named after the common wading bird (calidris alba). sandertoppene 72ø-107 (72°43.5´n 25°38.3´w). mountain sum mits in north lyell land south of kap alfred. named during the 1931–34 treårsekspeditionen by eugène wegmann as sander peaks, after bruno hermann max sander [1884–1979], a noted aus trian mineralogist and petrologist. sandgletscher 72ø-316 (72°11.8´n 25°50.8´w). glacier on the west side of schaffhauserdalen, with extensive sand and gravel moraines at its front. so named by john haller following explorations during lauge koch’s 1954 expedition, because he was stranded here with fritz schwarzenbach for three days during a violent sandstorm. sandodden 70ø-58 (70°33.0´n 25°51.3´w; map 4). pronounced sandy peninsula on the se coast of milne land. named by carl ryder’s 1891–92 expedition as sandodde. sandodden 74ø-96 (74°18.4´n 20°13.6´w). danish hunting station on the sw coast of wollaston forland on the north side of young sund. it was named after the sandy peninsula 4–5 km to the south, kap berghaus, which is known to norwegian hunters as heklas hvalrossnæs. the station (originally known as ny valdermarshaab) was built by østgrønlandske fangstkompagni in 1923 as a replacement for the station at kap borlase warren (valdermarshaab). the station was manned in the periods 1923–24, 1929–32, 1934–41, 1945–48 and 1949–50. it was taken over by nanok in 1929, and since 1952 has been used and maintained by sirius. it is said to be one of the best preserved of danish hunting stations (p.s. mik kelsen 1994). a danish hunter, axel kristensen, who died after being accidently shot in the arm at kap borlase warren in 1923, is buried here, as is eli knudsen, shot by german troops in 1943. the sirius headquarters, daneborg, are immediately adjacent to the station. sandstensdal 74ø-150 (74°24.8´n 20°15.7´w). valley in west wollaston forland, named during the 1931–34 treårsekspedi tionen by hans frebold. (sandstendal.) sandstensfjeldene 70ø-47 (70°43.0´n 25°22.3´w). range of hills on east milne land, nw of kap leslie, extending from kronen to hartz fjeld and southwards to slottet and glaukonitbjerg. named sandstens fjælde by carl ryder’s 1891–92 expedition, because of the abundant, light-coloured sandstones. (sandstensfjælde.) 296 sandstensodden 70ø (c. 70°39´n 25°17´w). name used in the 1891– 92 diaries of helge vedel (gulløv 1991) for the kap leslie area of east milne land. see also sandstensfjeldene. sandtorg 72ø (72°49.7´n 22°05.1´w). cape on east geographical society ø on the north side of cambridge bugt. used only on nsiu maps (lacmann 1937), the name was given for the locality of the same name in the troms district of norway. sandvik 74ø (74°09.2´n 21°31.4´w). norwegian hunting hut west of the mouth of granatdal, south clavering ø. it was built by the foldvik expedition in august 1926, and moved to this site in july 1927. the name appears on the nsiu (1932c) map and translates as ‘sandy bay’. the hut has also been known under the names granathytten, svampebugthytten, granitelva and stordal. sandøen 74ø-115 (74°15.8´n 20°09.4´w). small island in young sund, named by lauge koch’s 1929–30 expeditions as sand island, because it comprises exclusively sand and gravel. eiders and terns were reported to nest here in their thousands (pedersen 1960), and the locality was a bird sanctuary prior to establishment of the north–east greenland national park. walrus regularly come ashore here. (sandø, sandö.) sandøyra 72ø (72°58.3´n 22°13.9´w). flat, sandy delta on ne geographical society ø, so named on the nsiu maps of lacmann (1937). saníkivâjaq – see sanikkivaajak. sanikkivaajaq 70ø-363 (c. 70°29´n 21°58´w). coastal strip near to the town of scoresbysund. recorded by the 1955 geodætisk insti tut name registration, the name translates as ‘the poor side’. (saníkivâjaq.) sankt vitus bjerg 76ø-144 (76°38.0´n 25°09.5´w; map 4). moun tain on the south side of borgjøkel, dronning louise land. named by j.p. koch’s 1912–13 expedition in various forms (st. vitus bjærg, st. vitus-berg, st. vitus-spitze, st. vitus fjall.) sankt vitus is a catholic saint, whose memorial day (15 june) is the day the danish flag was said to have fallen from the sky at reval, estonia (see also revaltoppe and dannebrogsfjeldene). santes fair 71ø (71°54.5´n 24°43.5´w; map 5). mountain about 2100 m high on the west side of lower storgletscher, stauning alper. named by the 1961 bangor mountaineering club expedition. saqqaarissoq 70ø-146 (70°35.2´n 22°36.1´w). part of neill klinter between skævdal and astarte kløft, on the west side of hurry inlet. one of the names recorded by the 1955 geodætisk institut name registration, it roughly translates as ‘it has a pretty sunny side’. (sarqârigsoq.) sarpaq 69ø-52 (69°57.1´n 22°44.6´w). sound between an island and the coast ne of steward ø, north blosseville kyst, so named for the tidal current (= sarpaq). the name was recorded by the 1955 geodætisk institut name registration. sarpaq 72ø-279 (72°42.6´n 26°50.0´w). narrow channel at strøm næs, röhss fjord, marked by a strong tidal current. the green landic name, recorded by the 1955 geodætisk institut name registration, means ‘the current’. sarqârigsoq – see saqqaarissoq. satans galleri 72ø-507 (72°04.4´n 25°07.4´w; map 5). mountain ridge with a series of formidable pinnacles running nne of kors spids, south of gully gletscher. named by the 1963 cambridge university expedition saturn gletscher 71ø-336 (71°45.0n 24°53.8´w; map 5). glacier flowing south to join bjørnbo gletscher, south stauning alper. named saturn glacier by john hunt’s 1960 expedition, for the planet. sauruspasset 74ø-357 (74°35.7´n 20°19.2´w). pass at the south end of sillerendal, nw wollaston forland. named by wolf maync and andreas vischer during lauge koch’s 1938–39 expeditions for the important finds of fossil vertebrates. saussure massiv 73ø-324 (73°57.7´n 23°11.4´w; map 4). mountain in central hudson land. named by heinrich bütler during lauge koch’s 1938–38 expeditions after horace bénédict de saussure [1740–1799], a pioneer in the geography and geology of the alps. he had wide ranging scientific interests, discovered 15 new minerals, and encouraged the first ascent of mont blanc in 1786. he took part himself in the second ascent in 1787. (saussuremassiv.) saven 70ø-419 (70°40.7´n 29°35.5´w). nunatak group north of the upper part of rolige bræ. named by laurent jemelin during the 1967–72 ggu scoresby sund expeditions for a resemblance of the nunatak summits to the teeth of a saw. savkammen 71ø-266 (71°58.3´n 24°02.2´w; map 5). mountain ridge in the werner bjerge on the west side of kargletscher. named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk (savkammen = saw tooth comb). savoia halvø 70ø-362 (70°05.0´n 22°18.0´w; maps 3, 4). name used for the largely ice-free, triangular-shaped peninsula of which kap brewster is the ne point. the name was introduced by leonardo bonzi’s 1934 expedition as penisola savoia, who used it in a considerably wider sense than the present to include volquaart boon kyst and geikie plateau. it was named after the house of savoy, a historic dynasty of europe, and the ruling house of italy from 1861 to 1946. penisola italica has also been used. savryggen 76ø-337 (76°21.0´n 25°51.9´w; map 4). nunatak in sw dronning louise land, on the south side of budolfi isstrøm. so named by the 1952–54 british north greenland expedition be cause its profile of jagged peaks resembled the teeth of a saw. saxo bjerg 73ø-341 (73°19.4´n 22°20.3´w). mountain in the south giesecke bjerge. the name was proposed by the place name committee in 1939 to replace suggestions by wolf maync and andreas vischer. it commemorates the noted danish historian saxo [d. 1220], who wrote ‘gesto danorum’, a history of denmark in latin. skrukkryggen has also been used. (saxos bjerg.) scaphitesnæse 81ø (81°19.2´n 14°00.5´w). ridge in nw kilen, kron prins christian land, named after a cretaceous type fossil. the name is found on a coloured geological map of kilen printed in 1991 (pedersen 1991). schaffhauserdalen 72ø-117 (72°16.6´n 25°47.3´w; map 5). valley in ne nathorst land west of alpefjord. so named by eugène wegmann during the 1931–34 treårsekspeditionen, after the swiss town of schaffhausen. the valley is noted for widespread qua ternary moraines, which reminded wegmann of the old stony roads of schaffhausen. (schaffhausertal.) schalch bjerg 73ø-319 (73°52.5´n 23°25.6´w). mountain 1617 m high in central hudson land. named during lauge koch’s 1938– 38 expeditions by heinrich bütler after ferdinand schalch [1848– 1918], a german geologist noted for his work on the geology of baden and schaffhausen. (schalchs bjerg.) scheele bjerg 73ø-525 (73°08.0´n 25°56.7´w). mountain 1978 m high in ne suess land. named by a.g. nathorst in 1899 as scheeles berg, after carl wilhelm scheele [1742–1786], a swedish chemist noted in particular for his research in organic geochemistry. (scheele mountain.) schéele bjerg 72ø-184 (72°09.1´n 24°12.7´w; map 5). mountain between skeldal and store blydal, north scoresby land. named by prospecting teams associated with lauge koch’s 1948–49 expeditions after franz adolf von schéele, founder in 1830 of the noted engineering academy at filipstad. carl koch, the engineer responsible for establishing the mine at mestersvig, attended the academy. on some editions of the 1:50 000 scale topographic maps of mestersvig wittbergs bjerg is used for the same feature. (schéeles bjerg.) scheimpflugfjellet 72ø (72°55.0´n 22°36.8´w). mountain ridge in east geographical society ø, sw of laplace bjerg. used only on the nsiu maps of lacmann (1937), the name was given for theodor scheimpflug [1863–1911], an austrian who pioneered the practical use of aerial photography. 297 schiwiese 75ø (c. 75°19´n 17°50´w). feature in the vicinity of the base camp of the 1943–44 operation bassgeiger at kap sussi, shan non. the name is recorded by olsen (1965). schjelderup-hytta 73ø (73°36.3´n 22°02.2´w). this name has been used for a hunting hut in badlanddal near ladder bjerg, one of three huts built by ludolf schjelderup during the 1936–37 quest expedition. now a ruin. it has also been known as quest-hytten and tyvholmen. see also kapp schjelderup. (skeldruphytten, sjelderup hytten.) schnauder ø 78ø-12 (78°47.9´n 19°29.3´w; maps 1, 4). island in the north part of jøkelbugten, named by the 1906–08 danmarkekspeditionen as schnauder ö. koch (1916) writes that the 1906– 08 danmark-ekspeditionen owes a great debt of gratitude to professor schnauder. max schnauder [1860–1939] was a german astronomer and professor at the geodetic institute in potsdam, who had instructed alfred wegener in surveying observations and calculations (j. løve, personal communication 2009). (schnauders ø, schnauder island.) schneekoppe 75ø-17 (75°35.3´n 20°23.1´w). mountain 1417 m high in the north barth bjerge. so named by karl koldewey’s 1869–70 expedition, probably for its snowy summit and for a likeness to the mountain of the same name in the prussian riesen gebirge. (snetoppe.) schneekuppe 71ø (71°50.8´n 25°38.0´w; map 5). mountain between spærregletscher and the head of prinsessegletscher. it was named and climbed by the 1967 berchtesgaden expedition. it may be identical with berliner bjerg. scholanderdalen 72ø (72°53.6´n 23°26.6´w). valley on central geo graphical society ø draining south into vega sund. the name is used only on nsiu maps (lacmann 1937), and was given for per fredrik scholander [1905–1980], a norwegian artist and botanist who participated in nsiu expeditions to greenland, and subsequently carried out notable research at the naval arctic research laboratory in barrow, alaska. schrattenberg 73ø (73°25.6´n 26°29.1´w). ice dome about 2073 m high in southern andrée land. the name is found on a panorama drawn by john haller in 1949, reproducerd in schwarzenbach (1993). the word refers to the groved erosion features of karst limestones, and may here refer to similar textures in ice. schröter bjerge 71ø-174 (71°43.9´n 23°07.5´w). mountain range nw of fleming fjord. named by hans stauber during lauge koch’s 1938–38 expeditions after carl schröter [1855–1925], a swiss botanist and palaeobotanist and a specialist in alpine flora. schuchert dal 71ø-379a (71°30.0´n 24°24.0´w; map 4). large valley on the east side of the south stauning alper in which schu chert flod flows. the name was suggested by n.p. lasca following his work in 1966–67. schuchert flod 71ø-39 (71°30.0´n 24°24.0´w; maps 3–5). major n–s-flowing braided river at the east margin of the stauning alper, draining south into nordostbugt. mapped by lauge koch during flights in 1932 on the 1931–34 treårsekspeditionen expedition, and named after charles schuchert [1858–1942], an american geo logist and palaeontologist noted for his works on brachiopods. he had supplied koch with information on north american paleogeography. (schuchert river, schucherts flod, schuchertriver.) schuchert gletscher 71ø-156 (71°58.0´n 24°20.0´w; map 4). long glacier at the west margin of the werner bjerge, flowing from the central region of the highest mountains sw and south to schuchert dal. this name had originally been used by hans stauber in 1937, but the place name committee replaced it by the rarely used name kongespejlet. the latter was officially approved until 1956, when the name schuchert gletscher was revived at the suggestion of peter bearth, supported by lauge koch. schuchert–gully col 72ø (72°04.7´n 23°51.6´w; map 5). col be tween schuchert gletscher and gully gletscher. the name is used by bennet (1972). schuchert/örsteds pass 71ø (71°34.3´n 24°07.5´w). name used by hall (1964) for the pass between ørsted dal and schuchert dal. schultzhytten 77ø (c. 77°01´n 20°01´w). danish hunting hut built for nanok in the spring of 1938 east of trekronen, germania land, by carl henrik schultz. it has also been known as trekronerhytten, pashytten, hvalsletten and slettehytten. schwaben gletscher 71ø (71°46.9´n 25°39.1´w; map 5). glacier in the ne part of the borgbjerg gletscher region, southern stauning alper, nw of schwabentinde. probably named by the 1977 schwä bische stauning alper expedition. schwabentinde 71ø (71°46.2.4´n 25°39.1´w; map 5). peak 2376 m high in the ne part of the borgbjerg gletscher region, southern stauning alper. probably first climbed and named by the 1977 schwäbische stauning alper expedition. schwarck-tal 75ø (75°58.5´n 22°18.5´w). name used by curt teichert in 1932 during a journey along the margin of the inland ice. the four-man group descended from the ice along this valley, that corresponds to the position of lillegletscher and skyggesø south of the head of bessel fjord; the valley was named after a. schwarck, an assistant to the surveying parties in 1931 and 1932. schwarze wand 74ø-28 (74°46.4´n 20°07.7´w; map 4). mountain 1130 m high on south kuhn ø. so named by karl koldewey’s 1869–70 expedition, because its summit is formed by horizontal masses of black dolerite. however, it may also be named after the mountain of the same name in austria. the first ascent was made by julius payer in may 1870. (mt schwarze wand, black wall.) schwarzer zwilling 72ø (72°09.2´n 25°17.8´w; map 5). peak 2100 m high on the south side of vikingbræ, north stauning alper, climbed by hermann huber’s 1968 expedition. (black twin.) schwidefskyfjellet 72ø (72°44.7´n 22°28.9´w; fig. 14). mountain on se geographical society ø, sw of freycinet bjerg. so named on the nsiu maps of lacmann (1937) after kurt schwi defsky [b. 1905], who assisted in the preparation of the maps. schöne aussicht 71ø (71°58.9´n 25°33.4´w). peak 1640 m high on the east side of spærregletscher. this appears to be an alternative name for piz vadian, that was given when climbed by the 1966 berchtesgaden expedition. science valley 73ø (73°28.7´n 25°58.5´w). major e–w-trending valley in southern louise boyd land, so named by the 1999 cam bridge northeast greenland expedition, who made geological observations here. it has also been referrred to as jættedal. scimitar ridge 73ø (73°18.8´n 27°17.9´w). name used by the 1972 university of dundee expedition for a crescent-shaped ridge with a summit snow field, sw of haredalen in ne frænkel land. scioragletscher 72ø-311 (72°03.1´n 25°59.0´w). glacier on the north side of furesø, nathorst land. named during lauge koch’s 1954–55 expeditions by hans zweifel for its resemblance to scio ra gletscher in southern switzerland. scioraspids 72ø-479 (72°03.5´n 26°01.5´w). mountain on the north side of furesø, west of scioragletscher, nathorst land. named during the 1954–55 lauge koch expeditions by hans zweifel, after scioragletscher. scoop mountain 72ø (72°48.0´n 27°27.1´w). name used in the 1930s by louise a. boyd for lugano bjerg in gletscherland. as viewed from bocksrietdalen across hisinger gletscher the summit has a concave shape filled by a summit ice cap. louise boyd also labled this peak as c. mountain. scoresby land 71ø-141 72ø-14 (72°00´n 24°30´w; maps 3, 4). land area bounded to the north by kong oscar fjord and alpe fjord, and to the south by scoresby sund and nordvestfjord. the west boundary runs from borgbjerg gletscher via prinsesse gletscher to alpefjord. the official usage defined in 1961 includes the stauning alper, jameson land and liverpool land, although it is usually used in a more restricted sense for the north extension of jameson land and the stauning alper. carl ryder placed the name in an unmapped region to the north of nordvestfjord and west of 298 the werner bjerge on his 1891–92 maps. the name may have been adopted from a danish chart dated 1881, where it is placed at approximately 72°n. the name had earlier appeared on an 1844 map by j.d. pentonville published in london against the region 70°–75°n, and also occurs on the coast profiles of the 1879 ingolf expedition south of kap brewster at about latitude 69°n. in all cases the name commemorates the discoveries of william scores by jr. [1789–1857], arctic whaler and scientist, who was the first to make charts of this part of the east greenland coast. (scoresbyland, scoresbys land, scoresbysund landet.) scoresby sund [kangertittivaq / kangerlussuaq] 70ø-258 (70°17.0´n 23°00.0´w; maps 3, 4). major fjord up to 40 km wide leading west and nw to an extensive fjord system. named by william scoresby jr. in 1822 as scoresby’s sound after his father, who he describes as the original discoverer, and the first navigator to enter it. william scoresby senior [1760–1829] was an arctic navigator and whaler, who started in the greenland whale fishery in 1785, became a commander in 1790, and sailed nearly every year to the fishery until 1823. between 1796 and 1816 he had obtained 2693 tons of oil, the highest return of any whaling master. he is said to have invented the crow’s nest. scoresby sund is possibly the óllumlengri of the icelandic sagas (fig. frontispiece), the ‘fjord longer than all other fjords’. volquaart boon reported being carried into a large fjord by a current at about this latitude in 1761 when on a dutch boat – see also volquaart boon kyst. (scoresby-sund, scoresbysund fjord, scoresby’s sund, scoresby fjord, skoresbysund.) scoresby sund arkipelag 70ø, 71ø (70°–72°n). this was one of the physiographic divisions of east greenland proposed by storgaard (1927), and extended between latitudes 70° and 72°n, excluding jameson land and liverpool land. scoresbysund [illoqqortoormiut] 70ø-306 (70°29.1´n 21°57.9´w; maps 3, 4). town in south liverpool land in the ne part of rosen vinge bugt, north of the mouth of scoresby sund. the first houses were built by the expedition that founded scoresbysund in 1924– 1925, and settled by a group of about 70 greenlanders from ammassalik in 1925. the colony manager and the priest originally lived in one large house here, with the greenlanders mainly in the outlying settlements. a radio station and seismological station was built in 1927, a church and 10 houses in 1927–28, and a hospital was established in the french expedition house built in 1931, replaced by a new hospital in 1957. the seismological station was moved to kap tobin in 1963. the population of scoresbysund / illoqqortoormiut (ittorqqortoormiit) was 384 in 1986, with an additional 71 in outlying villages within the municipality boundaries. in 2007 the population was 529, with no permanent residents in the former settlements. scoresbysundvarden 80ø (80°34.5´n 18°26.5´w). cairn at the mouth of vardedalen on the the north side of ingolf fjord. the name was given by elmar drastrup’s 1938–39 expedition for the scoresby sund committee, which had given support to his expedition. the cairn was not observed by a geological party that camped at this site in 1995. scorpio 71ø (71°41.0´n 25°26.9´w; map 5). mountain about 2302 m high west of the head of jupiter gletscher, southern stauning alper. named and first climbed by the 1975 scottish scoresby land expedition led by e.a.m. walker for the constellation of the zodiac. scorpio glacier 71ø (71°56.3´n 25°26.6´w; map 5). glacier in the stauning alper flowing west to join duart gletscher at its confluence with spærregletscher. named by james clarkson’s 1961 expedition for scorpius, a constellation of the zodiac. scotstounhill 74ø-89 (74°12.4´n 22°36.5´w; map 4). large nuna tak 1254 m high in wordie gletscher nw of jordanhill. named by james wordie in 1926 for the scottish locality, which is spelt ‘scotston hill’ on modern maps. scott bjerg 73ø-724 (73°13.5´n 24°46.5´w). mountain massif 1723 m high on ymer ø, south of dusén fjord. named by peter friend during his 1968–70 expeditions after scott polar research insti tute, cambridge, at which he was based. the institute was founded in 1920 with the balance of proceeds of the public collections made following the deaths of robert falcon scott and four companions on their return from the south pole in 1912. scott keltie øer 72ø-62 (72°45.8´n 22°50.9´w; fig. 12). island group in vega sund. named by a.g. nathorst in 1899 as scott kel ties öar after john scott keltie [1840–1927]. keltie was secretary of the royal geographical society 1892–1915, and its increasing prestige in this period was said to be in large part due to his interests and influence. the name is currently applied to the small islands east of kista ø, but was apparently originally intended to include the larger islands of gåseøen, kista ø, magga ø and silja ø. (scott keltie islands, scott keltie-öyane.) scott’s inlet 73ø 74ø (74°05.0´n 19°53.0´w). this was the name used by william scoresby jr. in 1822 for gael hamke bugt, and it featured for a short period (1872–75) on british admiralty charts. it was named in honour of sir walter scott [1771–1832], scottish historical novelist and poet, perhaps most noted for his ‘waverley novels’. (scot’s inlet, scott’s einbucht, scotts indløb.) scout centenary 71ø (71°13.3´n 26°15.0´w). summit 2016 m high on the north side of edward bailey gletscher, renland. climbed and named by the 2007 west lancashire mountaineering group expedition. seanearbheinn 71ø (71°57.0´n 25°00.6´w; map 5). peak 2350 m high in the upper reaches of sefström gletscher, stauning alper. climbed by the 1998 scottish mountaineering club expedition. sechsspitze 74ø (74°35.4´n 19°08.8´w). name used by danish hunters for kronebjerg on west sabine ø, because of the many pinnacles which make up the summit (sechs = six). sevenspits has also been recorded. sederholm bjerg 73ø-80 (73°31.3´n 23°18.0´w; map 4). mountain on gauss halvø. named by helge g. backlund during lauge koch’s 1929 expedition in the form mt. sederholm to commemorate the authority on fennoscandian structures, jakob johannes sederholm [1863–1934]. a finnish petrologist and structural geologist, sederholm was director of the finnish geological com mission from 1892 to 1933. the mountain was climbed by group including backlund in august 1930. (sederholms bjerg, seder holm-berge.) sedgwick gletscher 72ø-242 (72°18.5´n 25°07.2´w; maps 4, 5). glacier in the north stauning alper, dividing the murchison bjerge. named during lauge koch’s 1950–51 expeditions by erd hart fränkl, after the noted british geologist adam sedgwick [1785–1873] who first applied the name cambrian to the geological period. sedimentkløft 77ø-26 (77°29.5´n 21°34.4´w). small ravine at the nw end of annekssøen, east of kofoed-hansen bræ. so named by the 1906–08 danmark-ekspeditionen, because of an outcrop of sedimentary rocks. (sedimentklöft.) seebach bjerg 75ø-7 (75°47´n 19°43´w; map 4). mountain 677 m high nw of roseneathbugt. named by karl koldewey’s 1869–70 expedition as cap seebach, after karl albert ludvig von seebach [1839–1880], a german geologist and palaeontologist who was pro fessor at göttingen. seebach had worked on rock samples brought back by koldewey’s first polar expedition (j. løve, personal communication 2010). as there is no real cape in the position indicated on koldewey’s maps the name was applied to the mountain forming the ‘cape’. a norwegian hunting hut was built east of the mountain by the 1932–34 giæver expedition. (sebachs bjerg.) seejoch 72ø (72°49.5´n 22°22.6´w). lake in adam af breemen dal, east geographical society ø, which drains into cambridge bugt. the name was used on a map by stauber (1938) describing work during lauge koch’s 1938–38 expeditions. seenplatte 71ø (71°45.9´n 27°07.6´w). name used by eduard wenk 299 for the plateau area north of central nordvestfjord, which has many lakes. one of the largest lakes is known as t-sø. seerakajik 70ø-278 (70°26.1´n 22°45.2´w). cove or lagoon in southern jameson land, dry at low water. recorded by the 1955 geo dætisk institut name registration, the name means ‘the little sandy beach’. (sêrakajik.) seerakajik 70ø-300 (70°30.0´n 22°05.4´w). river delta on the nw side of hvalrosbugt, south liverpool land. recorded by the 1955 geodætisk institut name registration, the name means ‘the little sandy beach’. (sêrakajik.) sefström gletscher 71ø-143 72ø-26 (72°03.0´n 25°20.0´w; maps 4, 5; fig. 77). glacier draining west from the stauning alper into alpefjord, where it merges with gully gletscher and almost blocks the fjord except for a narrow passage on the west side. named by a.g. nathorst in 1899 probably after nils gabriel sefström [1787– 1845], a swedish chemist and geologist noted for his discovery of vanadium, and investigations of glacial striae. (sefströms glacier, sefstrøm gletscher.) sefström tinde 72ø-269 (72°02.6´n 25°11.8´w; map 5; figs 27, 77). mountain 2714 m high on the ne side of sefström gletscher. the name seems to have been used first by hans gsellman’s 1957 expedition (which made the first ascent), and was approved in 1955 at the suggestion of john haller. the second ascent was made by the 1964 aac zürich expedition. (sefstroms tinde.) sefströmsgipfel 71ø (71°56.0´n 25°02.2´w; map 5). mountain 2622 m high at the head of the se branch of sefström gletscher. the name was used by hans gsellman’s 1957 expedition, which made the first ascent. sefstroms-lang col 71ø (71°56.8´n 25°01.4´w; map 5). col at the head of sefström gletscher leading to a tributary of storgletscher (formerly langgletscher). the name is used by bennet (1972). segelsällskapet fjord 72ø-31 (72°26.6´n 25°00.0´w; maps 4, 5). fjord between lyell land and the northern stauning alper, named by a.g. nathorst in 1899 after the kongelige svenska segelsäll skapet. (segelsälskapets fjord, royal yacht club fjord, yacht club fjord.) segldal 72ø-384 (72°05.7´n 23°31.5´w; map 5). valley between mesters vig and antarctic havn, north scoresby land. named by hans kapp during lauge koch’s 1957–58 expeditions. a norwe gian hunting hut near pictet bjerg sometimes goes under the name segldalen (see jostein). seglpas 72ø-383 (72°02.8´n 23°25.0´w). pass between the head of segldal and majdal, north scoresby land. named by hans kapp during lauge koch’s 1957–58 expeditions. seismisk station 71ø-309 (70°29.0´n 21°55.9´w). official designation for the seismic station at scoresbysund, erected by janus sør ensen in 1927–28. it was later moved to kap tobin. sejerstedt bødtkers hytta – see bødtkers hytta. seksogtredivekilometernæsset – see fyrretyvekilometernæsset. selwyn fjeld 72ø-500 (72°07.6´n 25°16.3´w; map 5). peak 2140 m high on the north side of gully gletscher. it was climbed by a cam bridge university expedition on 22 august 1963, and named after selwyn college, cambridge, incorporated into the university in 1882. (selwyn.) sem-gletscher 72ø (72°00.4´n 24°07.5´w; map 5). southern of three small glaciers between vestre gletscher and mellem gletscher in the north werner bjerge. the name was used by styger (1951) in his report on a climbing excursion during lauge koch’s 1950 expedition, and was named after shem (sem), the oldest son of noah. see also ham-gletscher and joffert-gletscher. semspitze 72ø (72°00.4´n 24°07.5´w). name used by styger (1951) for a mountain in the north werner bjerge at the head of semgletscher, between vestre gletscher and mellem gletscher. it was climbed by gerold styger with peter bearth during lauge koch’s 1950 expedition. see also sem-gletscher. sendlinger spids 71ø (71°52.9´n 25°25.5´w; map 5). mountain about 2300 m high between the upper part of duart gletscher and the upper basin of spærregletscher. climbed by karl herligkoffer’s expedition on 20 august 1966, and named apparently for send linger tor, one of the four town gates in central munich erected in 1318. (sendlinger bjerg.) sendlinger kalotte 71ø (71°53.1´n 25°26.7´w; map 5). mountain about 2250 m high on the ridge between duart gletscher and the upper basin of spærregletscher. climbed by karl herligkoffer’s expedition on 22 august 1966. sengstacke bugt 75ø-24 (75°20.8´n 18°15.8´w; map 4). bay on the north side of shannon. named by karl koldewey’s 1869–70 expedition as sengstacke bai, after heinrich sengstacke, 1st officer on the expedition ship germania. (sengstackes bugt, north bay.) sentinel 71ø (71°45.5´n 25°14.2´w; map 5). prominent mountain 2277 m high on the ne side of pegasus gletscher, south stauning alper. first climbed by james clarkson’s 1961 expedition, and so named because it appeared to guard the route to the upper reaches of bjørnbo gletscher. sentralen 72ø (72°23.1´n 25°15.1´w). original name for the hut now generally known as kap mæchel hytte. it was so named by fig. 77. the mountain sefström tinde (2714 m high) in the central stauning alper. the john haller photograph collection, geus archive. sefström tinde 300 peder sulebak, who helped build the hut for the møre expedition in august 1930. septembersø 72ø-276 (72°50.3´n 24°59.9´w). lake on central ella ø. so named by john w. cowie during work carried out during lauge koch’s 1952 expedition, because fossils were collected here in september. sêrakajik – see seerakajik. sermeq peqippaq 70ø (70°07.6´n 26°56.7´w). minor glacier on the south side of gåsefjord, west of sydbræ. between 2001 and 2007 the glacier advanced by 2.8 km, a phenomenon described as a svalbard-type surge (jiskoot & juhlin 2009). sernander bjerg 73ø-64 (73°41.6´n 22°41.8´w). mountain about 1600 m high in hudson land. named by helge g. backlund during lauge koch’s 1929 expedition in the form sernander ridge in honour of the noted swedish geologist, rutger sernander [1866– 1944], an expert on the post-glacial climatic evolution of fenno scandia. (sernanderberg, sernanderfjellet, mt. sernander.) seven pillars of hell – see jættevæggen. sevenspits 74ø (74°35.4´n 19°08.8´w). name used by danish hun ters for kronebjerg on west sabine ø, because of the many pinnacles which make up the summit. sechsspitze has also been recorded (hvidberg 1932). seward gletscher 69ø-44 (69°14.0´n 31°08.0´w). ne–sw-trending glacier nw of lindberg fjelde. named by l.r. wager’s 1935– 36 expedition as sewards glacier, after albert charles seward [1863–1941], a noted botanist and geologist, and professor of botany at cambridge from 1906 to 36, who had greatly helped the expedition. seward nunatakker 69ø-43 (69°16.0´n 31°14.0´w). ne–swtrend ing range of nunataks between seward gletscher and seward plateau, named by l.r. wager’s 1935–36 expedition as sewards nunataks. see also seward gletscher. seward plateau 69ø-42 (69°18.0´n 31°30.0´w). ice plateau nw of seward nunatakker, named by l.r. wager’s 1935–36 expedition as sewards plateau. see also seward gletscher. sfinks 70ø-373 (71°20.8´n 29°33.7´w). mountain in south paul stern land. climbed by paul stern during lauge koch’s 1958 expedition, and named after its resemblance to the sfinx-grat above the railway station of the jungfraujoch, bernese oberland, switzer land (fritz schwarzenbach, personal communication 1996). sfinksen 73ø-693 (73°22.9´n 26°18.6´w). mountain 2349 m high in south andrée land, with paw-like glaciers on the flanks and a shape resembling the head of a sphinx. named during lauge koch’s 1949–51 expeditions by john haller. sfinxen 70ø-343 (70°03.0´n 22°32.2´w). mountain 1268 m high on volquaart boon kyst, so named during the 1931–34 treårseks pe ditionen by laurits bruhn for a supposed resemblance to a sphinx. sfinxgletscher 70ø-344 (70°03.0´n 22°28.0´w). glacier east of sfinxen on volquaart boon kyst, so named during the 1931–34 treårsekspeditionen by laurits bruhn. shackleton bjerg 72ø-416 (72°53.8´n 28°46.6´w; map 4). pro m inent peak about 2900 m high in sw goodenough land. the name was given by james wordie in 1926, to commemorate sir ernest henry shackleton [1874–1923], a noted british antarctic explorer. shackleton was a member of scott’s 1901–04 antarctic expedition, led his own expedition in 1907–09 during which a new farthest south was reached, and also led the epic 1914–17 trans-antarctic expedition, when the endurance was trapped in the ice and sank. shackleton died on his way to the antarctic in 1923 and is buried in south georgia. wordie was one of the members of shackleton’s 1914–17 expedition. the peak was climbed by john haller on 2 august 1953. (schackleton bjerg, mount shackleton, shackleton moun tain.) shale peak 75ø (75°25.1´n 20°57.0´w). mountain north of arden caple fjord, nørlund land, climbed by michael banks and richard brooke in 1952 during the 1952–54 british north greenland expedition. it was named for the shales forming the summit (banks 1955). shangri-la 77ø (c. 77°00´n 24°46´w). sheltered corner of thomsen klippe or newton klippe, west dronning louise land. the locality was used as a camp site by the 1952–54 british north greenland expedition, and so named informally because it provided a haven from the piercing wind. shangri-la is the fictional monastry in james hilton’s novel ‘lost horizon’. shannon 74ø-12 75ø-50a (75°08.0´n 18°25.0´w; maps 2, 4). large island east of hochstetter forland. named shannon island by douglas clavering in 1823 after the royal navy frigate shannon, a 38-gun frigate on which he served as midshipman under sir philip broke. see also kap philip broke. ejnar mikkelsen occasionally used nordlandet and sydland for the northern and southern parts of shannon (j. løve, personal communication 2009). (shannon insel, shannonön, shannon ø, cannon-øya, sjannøy.) shannon sund 75ø-50 (75°12.4´n 19°09.4´w; map 4). broad sound between shannon and hochstetter forland. the name was said to have been used by danish hunters from about 1929, and first appears on the maps of the 1932 gefion expedition (jennov 1935). (shannonsund.) shannonfjorden 75ø (c. 75°05.0´n 19°44.0´w). name occasionally used by norwegian hunters and sealers for the fjord complex west of shannon, comprising the present shannon sund, peters bugt and ardencaple fjord (isachsen & isachsen 1932). sharks fin 70ø (70°47.0´n 26°16.1´w). culmination of a narrow ridge on the south side of korridoren, milne land. climbed, apart from the final 70 m, by the 2004 west lancashire scouts expedition. shell iskappe 76ø-317 (76°51.3´n 24°30.0´w; map 4). small ice cap in central dronning louise land, sw of army iskappe. named by the 1952–54 british north greenland expedition after the shell petroleum company, one of the two financial supporters of the expedition, which provided fuel, advice, facilities, and the loan of two seamen from their merchant navy fleet. shirley’s peak 72ø (72°06.6´n 24°55.5´w; map 5). peak in the stau ning alper on the ridge south of major passet. climbed by the 1996 scottish mountaineering club expedition. shrivenham 71ø (71°59.7´n 24°33.5´w; map 5). peak on the south side of schuchert gletscher about 1951 m high. it was climbed by the 1990 exercise green ice expedition. (mt. shrivenham). sidney fjeld 71ø-359 (71°58.0´n 25°07.5´w; map 5). peak 2300 m high overlooking the junction of canta bræ and sefström glet scher, stauning alper. climbed by a cambridge university expedition on 3 august 1963, and named after sidney sussex college, cambridge. see also sussex fjeld. sidselsøen 73ø-583 (73°58.8´n 24°15.1´w; map 4). small lake in south ole rømer land, named by sigurd skaun and harald welde in 1932 as sidseltjern. siegburger dal 72ø (72°05.3´n 23°58.0´w). name used by bierther (1941) for deltadal, the valley at the head of mesters vig, north scoresby land. it derives from work during lauge koch’s 1936–38 expeditions. (siborgdal, siegburger tal.) siestadal 72ø-474 (72°10.3´n 26°36.9´w). valley on the ne side of violingletscher, nathorst land. named during lauge koch’s 1954– 55 expeditions by hans zweifel. siestapasset 73ø-666 (73°41.0´n 25°32.9´w). pass between endeløs and spaltegletscher, west of randbjerg, andrée land. so named by erdhart fränkl during lauge koch’s 1948–50 expeditions because a rest was made here during a long traverse between the two glaciers. (siesta passet.) sigfriedbreen 74ø (74°22.8´n 21°06.7´w). glacier on north clav ering ø. so named on nsiu maps of lacmann (1937) after siegfried, hero of the german epic poem from about 1200, the nibelungenlied. (siegried glacier.) signes fjord 75ø (76°00.8´n 24°54.0´w). name used for the inner 301 branch of bessel fjord by poulsen (1991). sigurdsheim 74ø (74°50.5´n 19°45.3´w). norwegian hunting station on the east coast of kuhn ø, 3 km south of kap maurer. named after sigurd tolløfsen, whose expedition built the station in july 1932. siksakbjerg 73ø-108 (73°10.1´n 23°25.8´w). mountain 1084 m high on east ymer ø, so named during the 1931–34 treårseks peditionen by gunnar säve-söderbergh as mt. zigzag, because of the angular folding in the rocks. silberspitzen 71ø (71°53.9´n 25°34.8´w; map 5). name used by the 1964 aac zürich expedition for the peaks about 2400 m high on the ridge west of spærregletscher (silber = silver). the second ascent was made by karl herligkoffer’s 1966 expedition. the south ernmost peak is also known as breslauer spids. (silver peaks.) silja ø 72ø-333 (72°42.3´n 22°46.3´w; map 4). small island in vega sund. the name was proposed by søkortarkivet in 1956–57 fol lowing surveying of the channel through vega sund as an alternative approach for ships en route to mestersvig. it was given for the silja dan, a 4250 ton ice-strengthened polar ship built for the j. lauritzen shipping company in 1954 for the finnish trade. sold in 1964, it sailed as the veli until damaged by fire in 1971 and scrapped. silldal 71ø-414 (71°42.8´n 23°32.1´w). valley draining north into the upper part of ørsted dal. so named by katharina perchnielsen during the 1967–72 ggu scoresby sund expeditions because of the numerous dolerite sills. sillerendal 74ø-356 (74°38.0´n 20°18.4´w). valley in nw wol laston forland, so named during lauge koch’s 1936–38 expeditions by wolf maync, for the fossils. silvio bjerg 73ø-550 (73°05.3´n 27°54.0´w; map 4). mountain 2280 m high se of nordenskiöld gletscher, named by james wor die in 1929 as monte silvio. origin of name unknown. simonsen skær 76ø-83 (76°40.3´n 18°41.2´w). skerries on the east side of lille koldewey, named by the 1906–08 danmark-ekspedi tionen as simonsens skær. possibly named after the popular opera singer niels juel simonsen [1846–1906], as the expedition had a gramophone with them (jan løve, personal communication 2010). (simonsens skerries.) simpson dal 72ø-351 (72°08.7´n 22°11.0´w). valley on se traill ø, so named during lauge koch’s 1956–58 expeditions by h.p. heres. see also kap simpson. simpson-stranda 74ø (c. 72°07´n 22°15´w). name given to an intended norwegian hunting hut at kap simpson, se traill ø. material for the hut was deposited here by arktisk næringsdrift in 1929. a hut built here for sirius in 1955–56 is usually known under the name kap simpson hytten. sindalen 73ø-345 (73°22.8´n 22°09.1´w). valley in the southern giesecke bjerge, draining eastwards. the name was proposed by the place name committee in 1939 to replace suggestions by wolf maync and andreas vischer. it was probably given for the town of sindal in north jylland. stubbdalen has also been used. sinus gletscher 72ø (c. 72°95´n 25°06´w). minor glacier on the south side of gully gletscher, north stauning alper. named by sandro pucci’s 1984 expedition. sirius bjerg 71ø-263 (71°56.7´n 24°11.9´w). mountain 1632 m high in the werner bjerge north of sirius gletscher. named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk. sirius, the dog-star, is the brightest star in the northern night sky. sirius dal 79ø-23 (79°46.8´n 19°23.0´w; map 4). valley on sw hovgaard ø, providing a route from the front of nioghalv fjerdsfjorden to inner dijmphna sund. the name was approved in 1958, and was given for the sirius sledge patrol. sirius daneborg – see daneborg. sirius gletscher 71ø-270a (71°55.3´n 24°09.3´w; map 4). glacier in the werner bjerge, draining west to join schuchert gletscher. the name first appeared on the maps of styger (1951), in his de scription of a climbing excursion during lauge koch’s 1950 expedition, and was given for the star sirius. sista nålbrevet 71ø (71°02.2´n 25°29.2´w). name used by helge g. backlund during the 1931–34 treårsekspeditionen for the se pin nacle of one of the bjørneøer (island ix – see also bjørneøer), which was climbed in 1933 and used as a surveying point. siste-huset 71ø (71°38.0´n 22°23.7´w). norwegian hunting hut built by helge instad’s expedition in august 1932 at the head of nathorst fjord. it has also been known as bunn-huset. sivbreen 74ø (74°19.3´n 20°55.1´w). small glacier on central clavering ø. so named on nsiu maps of lacmann (1937) after sif, wife of tor in old nordic mythology. sjapgletscher 72ø-156 (72°18.0´n 22°37.9´w). glacier on se traill ø, south of mountnorris fjord. so named during lauge koch’s 1938–38 expeditions by hans p. schaub because it is covered in slush and water during the summer (sjap = slush). sjelnan 76ø (76°15.8´n 21°41.4´w). name sometimes used for the norwegian hunting hut at kap ullidtz, built in august 1933 for john giæver’s expedition. sjoaelv 73ø-169 (73°28.0´n 21°14.3´w). river on the south side of hold with hope, named on an nsiu map (1932a) in the form sjoa (fig. 13), possibly for a river of the same name in the oppland area of norway. the norwegian word implies a singing noise, often of a river. sjussen 71ø-436 (71°11.5´n 28°28.1´w; map 4). ice-dammed lake between the front of vindue gletscher and eielson gletscher, which periodically drains to leave a chaos of stranded icebergs. named by johan d. friderichsen during the 1967–72 ggu scores by sund expeditions after ‘sjus’, a slang expression for whisky and soda with ice. sjællandselv 70ø-98 (70°40.2´n 23°36.4´w; map 4). river in south jameson land flowing sw into scoresby sund south of vandre blokken. named during the 1931–34 treårsekspeditionen by lau rits bruhn for the island of sjælland, denmark. skalbæk 81ø (81°18.1´n 13°44.0´w). stream in nw kilen, kron prins christian land, where fossil mussels and ammonites are com mon in sandstone concretions. the name is found on a coloured geological map of kilen printed in 1991 (pedersen 1991). skallingen 79ø-38 (79°50.0´n 22°00.0´w; maps 1, 4). large area of south kronprins christian land, limited to the west and north by græsdal and sæfaxi elv. mapped by lauge koch during flights in 1933 on the 1931–34 treårsekspeditionen, and named after the danish locality of the same name near esbjerg. skandalen 73ø (73°33.3´n 20°30.5´w). norwegian hunting hut on the east coast of hold with hope, nw of kap broer ruys, built by the foldvik expedition in august 1927. the name occurs in the list of huts by orvin (1930). bukta, tvivlsom and moskusoksehytta have been used for the same hut. skansekløft 76ø-304 (76°57.4´n 20°04.7´w). depression north of østre skanse and vestre skanse, south germania land. so named by the 1938–39 mørkefjord expedition. skansen 71ø-43 (71°09.4´n 22°41.7´w; map 4). hill 690 m high sw of carlsberg fjord. named by g.c. amdrup’s 1898–1900 expedition. skardvatnet 72ø (72°52.9´n 22°29.7´w). lake on east geographical society ø, on the north flank of leitch bjerg. used only on nsiu maps (lacmann 1937), and so named because it lies on a pass (= skard). skarndal 71ø-134 (71°04.5´n 22°15.4´w). valley on the north side of the head of storefjord, central liverpool land. so named by helge g. backlund during the 1931–34 treårsekspeditionen, for the occurrence of skarn minerals. skarren 73ø-335 (73°32.4´n 22°09.2´w). mountain 1150 m high in the north giesecke bjerge. this may have been adopted from the form skarven used on an nsiu map (1932a), possibley given after 302 one of several similar place names in norway. de saussure bjerg has also been used. skartind 72ø (72°03.7´n 24°54.2´w). snow summit about 2310 m high on the east side of crescent col at the head of gully gletscher, stauning alper. climbed and so named by the 1996 norwegian stau ning alper expedition. skaunhøgda 72ø (72°48.4´n 22°14.6´w). plateau on east geo graphical society ø, on the ne flank of freycinet bjerg. used on the nsiu maps of lacmann (1937), the name was given for sigurd skaun [b. 1894], a norwegian journalist who accompanied the 1932 nsiu expedition to east greenland. skeen 80ø (80°34.4´n 19°31.1´w). glacier on the west side of the prinsesse caroline-mathilde alper, inner ingolf fjord, named by elmar drastrup’s 1938–39 expedition for its spoon-like shape. the name is also found on 1957 ams maps. skelbræ 72ø-526 (72°05.0´n 24°24.0´w; map 4). glacier at the head of skeldal leading to skelpas. the name was suggested by n.p. lasca following work in 1966–67. skeldal 72ø-99 (72°15.4´n 24°15.5´w; maps 4, 5). broad valley on the east flank of the north stauning alper, dividing the mountains to the west from the lower region to the east. named by ove simon sen during the 1931–34 treårsekspeditionen (skel = dividing line). skeldal elv 72ø-524 (72°15.1´n 24°14.2´w; map 5). river at the east margin of the stauning alper in the valley skeldal. the name was approved at the suggestion of n.p. lasca following his work in 1966–67, but had occasionally been used earlier in geological publications. (skel-fluss.) skeldal-hytta 72ø (72°17.5´n 24°08.9´w). name generally used for the norwegian hunting hut east of the mouth of skeldal, se of menander øer. originally known as elveidet, it was built in august 1930 for the møre expedition. (skelhytte, skeldalhyyen, skjell dalen.) skelelv 70ø (70°31´n 22°09´w). name used by rosenkrantz (1942) for a river in south liverpool land following the boundary be tween sedimentary and crystalline rocks. skelhøje 72ø-87 (72°32.3´n 22°59.1´w; map 4). hill about 500 m high on east traill ø, nw of mountnorris fjord. it was named during the 1931–34 treårsekspeditionen by ove simonsen for the danish locality of the same name in jylland. skelpas 72ø-298 (72°01.1´n 24°21.0´w; map 5). pass between skålen, a major branch of schuchert gletscher, and skelbræ, in the werner bjerge. the name first appeared on the maps of styger (1951), and derives from a climbing excursion during lauge koch’s 1950 expedition. (skel pass.) skibakken 70ø (c. 70°27´n 26°15´w). small isolated hill on danmark ø, probably situated just ne of hekla havn. the name is only used in the expedition report by hartz (1895) on work during carl ryder’s 1891–92 expedition. skibselv 72ø-222 (72°08.1´n 23°51.9´w; map 5). river draining the east side of blyryggen, which reaches mesters vig beside ekspedi tionshus. so named by prospecting teams associated with lauge koch’s 1948–49 expeditions, possibly because ships took on fresh water here. it corresponds to river 1ø on detailed maps of the region. (skibselven.) skibshavn 76ø (76°45.7´n 18°41.3´w). alternative name for danmark havn used by trolle (1913) in his reports on the 1906–08 danmark-ekspeditionen. skibssø 76ø-189 (76°46.5´n 18°42.6´w). lake due north of the anchorage of the expedition ship danmark in danmark havn, named during the 1906–08 danmark-ekspeditionen. the staff at danmarkshavn weather station use the name vandsø. skida 72ø (72°16.3´n 23°55.9´w). hut in a small bay nw of ny havn. the name seems also to have been used by the university of dundee expeditions between 1968 and 1974 for the bay. the hut was originally a bath house and toilet building at mestersvig air field, and was moved to this site by airfield personnel in 1965. skiferbjerg 72ø-264 (72°09.9´n 25°18.8´w; map 5). mountain 1970 m high in the north stauning alper south of vikingebræ. it was climbed by the dansk-norske ekspedition on 17 august 1954. the name was proposed by john haller who explored the same region during lauge koch’s 1954 expedition (skifer = slate, shale). skiferdal 74ø-152 (74°25.6´n 20°16.3´w). valley in west wollaston forland, so named by hans frebold during the 1931–34 treårseks pedi tionen for the shaly rocks (= skifer). skiferkløft 74ø-200 (74°15.7´n 20°25.8´w). ravine on ne clav ering ø, draining into young sund. the name is used by bøgvad & rosenkrantz (1934) in the form skifferkløften, and refers to the occurrence of shale. skildpadden 73ø-249 (73°05.0´n 22°29.7´w). island in the broch øer group. named on the 1932a nsiu map as skjelpadda, presumably for a resemblance to a turtle. skildvagten 73ø-527 (73°03.4´n 25°09.6´w). prominent pointed mountain 1046 m high in east suess land. named skiltvakten by a.g. nathorst in 1899 because it stood like a sentry (= skildvagt) at the entrance to the unexplored waters to the south. (the sentinel, mt skildvagten.) skillebugt 71ø-416 (71°14.0´n 25°41.7´w). deep bay in se renland, which divides (= skille) two areas with different geological structures. named during the 1967–72 ggu scoresby sund expeditions by niels henriksen. skilledal 74ø-317 (74°24.2´n 20°57.0´w). valley on north clav ering ø, which with skillegletscher divides clavering ø into two equal parts. the name is attributed to richard foster flint and arises from work during louise boyd’s 1937 expedition. nivlheimdalen has also been used. (skille valley.) skillegletscher 74ø-318 (74°16.9´n 20°56.9´w). glacier on north clavering ø, which with skilledal divides clavering ø into two parts. the name is attributed to work by richard foster flint during louise boyd’s 1937 expedition. vintergata has also been used. (skille glacier.) skillingen 72ø (72°49.6´n 22°56.9´w). island in vega sund, nw of gåseøen. so named on the nsiu maps of lacmann (1937). skinfaksebreen 74ø (74°22.0´n 20°41.8´w). small glacier on ne clavering ø. used only on nsiu maps (lacmann 1937), and named after skinfaxe (or skinfakse) of old nordic mythology, the horse of the day whose shining mane lights up the earth. skipperdal 72ø-245 (72°23.6´n 24°51.5´w; map 5). valley in the north stauning alper. the name was given by the place name committee as a substitute for jassdal, proposed by erdhart fränkl during lauge koch’s 1950–51 expedition. it may commemorate axel jensen, who was skipper of the polypen and assisted fränkl in 1950. see also akselborg. skjerva 73ø (73°23.5´n 23°08.4´w). river on south gauss halvø, flowing in elsa dal. so named on an nsiu map (1932a), perhaps after one of several localities of the same name in norway. skjervens tind 71ø (71°53.5´n 25°06.0´w). mountain about 2350 m high on the north side of roslin gletscher, between the two branches of the minor glacier valhallbreen. it was climbed by the 1996 norwegian stauning alper expedition, and so named after ove skjerven [1944–1984] a colleague who had died while climbing in peru. (skjervens topp.) skjoldet 71ø-288 (71°53.1´n 24°02.3´w; map 5). summit on the ridge between aldebaren gletscher and breithorn gletscher, south werner bjerge. named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk (skjoldet = the shield). skjoldgletscher 72ø-162 (72°33.3´n 22°29.8´w). glacier on se traill ø on the north side of mols bjerge. named during lauge koch’s 1936–38 expeditions by hans p. schaub for its shape. skjoldungebræ 72ø-97 (72°18.7´n 24°44.4´w; maps 4, 5; fig. 78). large glacier in the north stauning alper, draining north to kong oscar fjord. named during the 1931–34 treårs eks peditionen 303 after the large island skjoldungen in se greenland. skjærbosund 70ø (70°17.0´n 23°00.0´w). a variation of scoresby sund, occasionally used by norwegian hunters and sealers (e.g. isach sen & isachsen 1932). skogulfjellet 73ø (73°08.8´n 23°46.3´w). mountain on se ymer ø. so named on an nsiu map (1932a), possibly a derivation from the norwegian word for a wood or forest. skolma 73ø (73°22.1´n 23°04.0´w). stream on south gauss halvø, flowing in agda dal. so named on the 1932a nsiu map. skorfjellet 74ø (74°09.5´n 21°08.5´w). mountain ridge on south clavering ø, running from the present vesttinden to østtinden. so named on the nsiu maps of lacmann (1937) for the prominent gulleys (= skorn) cut by a series of streams. see also skårene. skorfjellfonna 74ø (74°11.0´n 21°10.2´w). small ice cap on south clavering ø, the present taggletscher. so named on the nsiu maps of lacmann (1937) for its proximity to the skorfjellet of nor wegian maps. skorpa 72ø (72°40.6´n 22°23.1´w; fig. 14). narrow elongate island east of nordenskiöld ø. used on the nsiu maps of lacmann (1937), the name was adopted from a locality of the same name in the troms district of norway. skottepasset 69ø-74 (69°49.5´n 23°39.0´w; map 4). pass between the head of trolddal and steno gletscher. named by malcolm slesser’s expedition in 1969, after the scottish member of the expedition (this was slesser himself ). skotsketinde 72ø (72°07.6´n 24°45.3´w; map 5). mounain 1775 m high on the west side of upper bersærkerbræ, northern stauning alper. skrubbfjellet 73ø (73°23.6´n 22°07.2´w). mountain ridge north of sindalen in the south giesecke bjerge. so named on an nsiu map (1932a), and probably derived from the norse word for wolf. skrubbtind 73ø (73°25.5´n 22°18.1´w). mountain in the south giesecke bjerge corresponding to the present svanning bjerg. so named on the 1932a nsiu map, and derived probably from the norwegian word for a wolf. skrukkedalen 73ø (73°17.9´n 22°13.1´w). valley south of skrukk ryg gen in the south giesecke bjerge, corresponding to the present vilddalen. so named on the 1932a nsiu map. see also skrukk ryg gen. skrukkryggen 73ø (73°19.4´n 22°20.3´w). ridge in the south gie secke bjerge, equivalent to saxo bjerg. so named on an nsiu map (1932a) for the wrinkled or puckered appearance of the ridge slopes due to the numerous minor drainage channels. skrællingedalen 74ø-250 (74°08.1´n 20°55.5´w). valley on south clavering ø reaching the coast at dødemandsbugten. there are three large inuit settlements east of the mouth of the river. the name was used first on the nsiu (1932a) map in the form skræl inge dalen, and derives from the old norwegian word for the inuit (= skrællinger). skrællingelven 74ø (74°08.1´n 20°55.5´w). name used by glob (1946) for the river in skrællingedalen, south clavering ø. skræntdal 72ø-430 (72°38.7´n 27°21.1´w). valley west of the head of röhss fjord, so named by ove simonsen during the 1931–34 treårsekspeditionen because of its steep sides. skrænterne 70ø-88 (70°04.4´n 24°35.1´w; map 4). range of cliffs set back from volquaart boon kyst, so named during the 1931–34 treårsekspeditionen by laurits bruhn (skrænterne = the cliffs). skrænthytten 74ø (74°28.7´n 21°53.5´w). name used for the norwe gian hunting hut 3 km nw of kap ehrenberg in tyrolerfjord, built for finn devold’s expedition in september 1928. it has also been known as tyrolerheimen. skråbræ 73ø-581 (74°02.2´n 28°50.0´w). glacier between hobbs land and arnold escher land, named by arne høygaard and mar tin mehren in 1931 as skråbreen because it descends steeply down to join adolf hoel gletscher (skrå = sloping, oblique). skuvfjellet 73ø (73°24.1´n 22°29.1´w). mountain on gauss halvø, corresponding to part of højsletten. so named on the 1932a nsiu map, the name in norwegian dialect implying something uplifted or prominent. skygge fjord 76ø (76°15.0´n 21°01.3´w). narrow fjord in north ad. s. jensen land, the present syttendemajfjord. one of the names on the 1932 edition of the geodætisk institut 1:1 million scale map, it derives from lauge koch’s aerial observations during the 1931–34 treårsekspeditionen (skygge = shadow). skyggedal 72ø-108 (72°45.8´n 25°31.9´w). valley in lyell land south of kap alfred. the name was given by the place name com mittee in 1935, and records that the valley is often in shadow (= skygge), being open only to the ne. skyggesø 75ø-85 (75°58.0´n 22°13.5´w; map 4). lake west of the head of bessel fjord. the name was suggested by the place name fig. 78. looking north-east from skjoldungebræ across kong oscar fjord to svinhufvud bjerge on traill ø. sorteelv gletscher and syltopperne are carved into the brightly coloured rocks of the eleonore bay supergroup. the john haller photograph collection, geus archive. kong oscar fjord syltopperne skjoldungebræ ssoorrtteeeellvv gglleettsscchheerr sorteelv gletscher traill ø svinhufvud bjerge 304 committee in 1935, and records that the lake is often in shadow. skylstad 75ø (75°14.9´n 20°52.8´w). norwegian hunting hut built in august 1932 for john giæver’s expedition on the south side of the mouth of kildedalen. giæver named it for his friend jakob skyl stad [b. 1888], editor of the trondheim newspaper ‘nasjonbladet’. skylstaddalen 73ø (73°00.0´n 23°30.8´w). valley on central geo graph ical society ø west of rudbeck bjerg, draining north into sofia sund. so named on the nsiu maps of lacmann (1937) after jacob skylstad – see skylstad. skærene 79ø (79°37.1´n 19°29.6´w). group of small islands off the front of nioghalvfjerdsfjorden also known as bloch nunatakker. the name was used by the 1996 mylius-erichsen’s minde ek s pedition. skærfjorden 77ø-35 (77°25.0´n 19°15.0´w; maps 2, 4). broad irregular fjord north of germania land with many islands and skerries, so named by the 1906–08 danmark-ekspeditionen. baie d’orleans has also been used. (reef fjord, skær fiord, skjærfjorden.) skævdal 70ø-143 (70°34.8´n 22°37.7´w). name proposed by alfred rosenkrantz for a small valley on the west side of hurry inlet (skæv = crooked). skævelv 70ø-143a (70°34.8´n 22°37.7´w). river flowing in skæv dal on the west side of hurry inlet. the name was suggested by alfred rosenkrantz. skåldal 70ø-176 (70°36.9´n 22°10.7´w). valley in south liverpool land draining nw into gubbedal. so named by laurits bruhn during the 1931–34 treårsekspeditionen for its shape (skål = bowl). skåldalen 72ø (72°59.0´n 24°11.5´w). valley on west geographical society ø draining north into sofia sund. used only on nsiu maps (lacmann 1937), and so named because it drains a bowlshaped corrie. skålen 72ø-299 (72°00.8´n 24°18.4´w; map 5). bowl-shaped glacier, an upper lobe of schuchert gletscher. the name was adapted from the grosses becken of styger (1951), a name used on maps of a climbing excursion to the region during lauge koch’s 1950 expedition. (skaalen.) skårene 74ø-253 (74°10.9´n 21°05.5´w). mountain on south clavering ø. adapted from the skorfjellet of norwegian maps used for a nearby mountain ridge (see skorfjellet). (skaarene.) skårkammen 71ø-280 (71°54.5´n 24°06.1´w; map 5). mountain ridge on the south side of fingerbøllet at the head of aldebaren gletscher, werner bjerge. named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk. skårungane 72ø (72°41.8´n 22°27.2´w; fig. 14). small island in east vega sund, north of nordenskiöld ø. the name was used only on nsiu maps (lacmann 1937), and was given for the young gulls. slambugten 76ø-274 (76°55.3´n 20°03.8´w). bay east of hvalros odden on the south coast of germania land. so named by the 1938–39 mørkefjord expedition, for its muddy water. slamodden 76ø-352 (76°55.1´n 19°56.8´w). peninsula east of slambugten. the name was reported as in general use by the staff at danmarkshavn weather station in the period 1969–71. slamsø 75ø-47 (75°15.4´n 21°42.7´w; map 4). lake in kildedalen, c.h. ostenfeld land. the name refers to the muddy waters of the lake, and is first found on the 1932 edition of the lauge koch’s 1:1 million scale map published by the geodætisk institut. slamsøen 77ø (77°08.6´n 23°24.6´w). name occasionally seen used for britannia sø in northern dronning louise land, and arising from the muddy waters. slanstinde 72ø (72°09.1´n 25°04.9´w; map 5). peak 2350 m high in the north stauning alper, on the south side of vikingebræ, climbed by claude rey’s 1970 expedition. exact position uncertain according to bennet (1972). slate river 70ø (70°31.5´n 22°48.8´w). name used by hermann aldinger during the 1931–34 treårsekspeditionen for the present ostreaelv, a river in south jameson land where there are slaty rocks (aldinger 1935). sletta 74ø (74°35.7´n 19°51.4´w). norwegian hunting hut on the east side of albrecht bugt, wollaston forland built by the hird expedition in 1928. named after the wide plain (= sletta) where it is situated. it has also been known as græstorvshytten. (sletten, slettehuset, slette huset.) slettedalen 74ø-292 (74°34.5´n 21°00.0´w; map 4). broad flat val ley between lindeman fjord and store sødal. the name is attributed to the wintering party at kulhus in 1935. slettefloden 74ø-352 (74°32.6´n 19°55.1´w). river draining stor sletten in north wollaston forland. named during lauge koch’s 1938–38 expeditions by wolf maync and andreas vischer. slettehuset – see sletta. slettehytten 74ø-f157 (74°35.4´n 20°02.9´w). danish hunting hut built for nanok in may 1947 on the west side of storsletten, wolla ston forland. slettehytten – see trekronerhytten. slien 71ø-110 (71°11.1´n 21°50.6´w). fjord in east liverpool land ne of kap james. named during the 1931–34 treårsekspeditionen by laurits bruhn. slippen 73ø-103 (73°20.0´n 23°50.1´w). valley on the north side of gunnar andersson land, ymer ø, draining ne. named by th. johansen during the 1931–34 treårsekspeditionen. it is a long narrow ravine with a moderate gradient (slippen = the slipway). slippenhytten 73ø (c. 73°21´n 23°46´w). norwegian hunting hut built in august 1938 for ole klokset’s expedition at the mouth of fig. 79. the nunatak slottet near eleonore sø, whose conspicuous summit is formed of yellowwhite cambrian quartzite. this formation exposed only in the western nunatak region of northern east greenland is the source of the widespread erratic blocks (dropped by glaciers) of skolithus quartzite. 305 the valley slippen, gunner andersson land. it has also been known as kloksethytten and kap martha hytten. slottet 70ø-44 (70°41.3´n 25°19.4´w). minor summit on a ridge nw of kap leslie, east milne land. named by hermann aldinger during the 1931–34 treårsekspeditionen in the form schloss (= castle = slot). (castle hill). slottet 73ø-412 (73°57´n 28°15´w; map 4; fig. 79). imposing nuna tak of white quartzite at the east end of langeryg, arnold escher land. named during lauge koch’s 1951 expedition by hans r. katz for its supposed resemblance to a castle. slottneset 72ø (72°53.1´n 21°54.5´w). cape on east geographical society ø, corresponding to the south flank of kap mackenzie. so named on the nsiu maps of lacmann (1937) because it resembles a castle (= slot) in shape. slugtdal 72ø-389 (72°01.4´n 23°17.6´w). valley west of antarctic havn, north scoresby land. so named by hans kapp during the 1957–58 lauge koch expeditions, because of the marked ravine in the valley. (slugtdalen.) slugtdalen 74ø-324 (74°02.1´n 22°52.8´w). valley in north hud son land draining north to wordie gletscher. named during lauge koch’s 1938–38 expeditions by heinrich bütler. mehren dalen is used on lacmann’s (1937) maps. slyngelv 76ø-306 (76°56.6´n 20°13.1´w). minor tributary to lakse elven ne of mørkefjord station. so named by the 1938–39 mørke fjord expedition for its strongly meandering course (slynge = swing). slædedalen 74ø (74°34.0´n 20°17.7´w). valley in nw wollaston forland south of sauruspasset, part of the present canyondalene. the name was used by wolf maync (1947) who made a sledge journey from kuhn ø to clavering ø following this route during lauge koch’s 1938–38 expeditions (slæde = sledge). slædelandet 77ø-109 (77°08´n 19°52´w; map 4). relatively lowlying region of germania land bounded by valdemarsmuren to the west, and moskusoksefjeldene to the se. so named by the 1938–39 mørkefjord expedition since all northward sledge journeys from mørkefjord station went this way, although the first 20 km was generally snow-free and gave poor sledging. slædepas 73ø (c. 73°44´n 20°30´w). name used by gelting (1937) for a locality in the vicinity of knudshoved, hold with hope; exact location uncertain. possibly a hunters name. slædepasset 76ø-281 (76°24.0´n 20°53.3´w). sledge route from sylbugten to the hunting station ålborghus. named during the 1938–39 mørkefjord expedition. slædeøen 76ø (76°14.5´n 19°52.6´w; map 4; fig. 80). island east of kap peschel, ne of ad. s. jensen land. nyholm-poulsen (1985) reported finding two of j.p. koch’s horse sledges on an iceberg in 1933, which had been abandoned on brede bræ in 1912. one of the sledges was dragged to the north point of the island on 28–29 june 1933, and now bears a memorial plaque to j.p. koch. smalle spærregletscher 80ø-113 (80°37.0´n 18°43.0´w; map 4). glacier draining south into ingolf fjord, opposite brede spærre gletscher. probably named by john haller following his explorations during lauge koch’s 1956–58 expeditions. smallefjord 75ø-37 (75°27.8´n 21°45.3´w; map 4; figs 51, 81). nar rower of the two branches of ardencaple fjord, named in this form by the 1906–08 danmark-ekspeditionen. (smalle fjord.) smallefjordhytten 75ø (75°27.8´n 21°38.5´w). norwegian hunting hut on the north side of smallefjord, built in august 1933 for john giæver’s expedition, and still standing in 1988. it has also been known as tornøestua. smallegletscher 72ø-319 (72°03.0´n 25°46.4´w). long, narrow glacier on the north side of eastern furesø. named by john haller following explorations during lauge koch’s 1954 expedition. smalleryg 72ø-438 (72°02.7´n 27°09.3´w; map 4). long and nar row ridge between jomfrudal and grænsedal, nathorst land, named by ove simonsen during the 1931–34 treårsekspeditionen (smalleryg = narrow ridge). smalleryg 74ø-216 (73°59.3´n 21°25.0´w). minor, narrow ridge on the ne slope of frebold bjerg, between river 13 and river 14, nw hold with hope. named by eigil nielsen during the 1931–34 treårsekspeditionen (teichert & kummel 1976). depot ryg (depot ridge) is in the same area). smedal 73ø (73°18.7´n 22°41.8´w). norwegian hunting hut on the south side of gauss halvø, west of kap franklin, built by john giæver and halvard devold for arktisk næringsdrift in august 1930. named after gustav smedal, a norwegian lawyer, chairman of norges grønlandslag and norges ishavsråd, and much con cerned with the conflict over norwegian rights in east greenland. it is also known as margrethedalshytten. (smedals hyt ten, smedalen, smedahl.) smedal valley 73ø (73°19.8´n 22°34.1´w). name occasionally used for the present margrethedal west of kap franklin, where the nor wegian hunting hut smedal is situated. smith woodward bjerg 73ø-110 (73°26.5´n 23°20.0´w). moun tain on the sw coast of gauss halvø, named during the 1931–34 treårsekspeditionen by gunnar säve-söderbergh as mt. smith woodward. sir arthur smith woodward [1864–1944] was a british vertebrate palaeontologist noted for his work at the british museum (natural history). he was the first to describe devonian vertebrate fossils from east greenland. norwegian maps of the 1930s use havgrimfjellet and einarfjellet for approximately the same mountain. (smith woodwardsberg.) smith’s island 71ø (71°44.5´n 22°14.1´w). supposedly an island adjacent to canning land, but probably the mountain behind the fig. 80. two of the horse-sledges of j.p. koch’s 1912–13 expedition abandoned on bredebræ in 1912 were discovered in 1933 on an iceberg in dove bugt by danish trappers. one of the sledges was dragged to an island off kap peschel, now known as slædeøen. 306 present kap tyrell (fig. 3). it was named by william scoresby jr. in 1822 for sir james edward smith [1759–1828], a botanist notable for his purchase of the entire library and collections of the younger linnæus. smith founded the linnean society in 1788, and was its first president. (smiths ö.) smøgen 75ø-65 (75°48.2´n 20°55.9´w). valley on the south side of langsø, nørlund land. the name originated from the wintering party at kulhus during the 1931–34 treårsekspeditionen (smøge = narrow passage, alley). småskærene 77ø-65 (77°26.1´n 19°37.0´w; map 4). group of small skerries se of joinville ø in skærfjorden, so named during the 1931–34 treårsekspeditionen by david malmquist. snaddheimen 73ø (73°10.2´n 26°40.0´w). norwegian hunting hut on the coast of east frænkel land, south of niggli dal. built by bjarne and oddvar akre for arktisk næringsdrift in august 1938, and named for the ringed seal (= snadd), which is very common in the fjords. reported as a ruin in 1976. snedrivegletscher 72ø-480 (72°18.1´n 26°06.9´w). broad glacier at the head of schaffhauserdal, nathorst land. named during lauge koch’s 1954–55 expeditions by hans zweifel (snedrive = snow drift). snefnugdal 70ø-450 (70°21.3´n 29°24.0´w). valley in se paul stern land draining into vestfjord gletscher. so named by w.e. adrien phillips during the 1967–72 ggu scoresby sund expeditions because he was snow-bound in camp here for four days in 1972 (sne fnug = a porous snowflake). snefogsdepotet – see fyrretyvekilometernæsset. snegryden 70ø-447 (70°20.6´n 29°58.8´w). nunatak in west paul stern land. named by w.e. adrien phillips during the 1967–72 ggu scoresby sund expeditions for the round cauldron-like depression in the centre of the nunatak which collects snow. sneharefjeld 73ø-419 (74°00.6´n 27°06.5´w). nunatak in northernmost andrée land, so named by hans r. katz during lauge koch’s 1951 expedition because of the sighting of a hare. snehorn 70ø-372 (71°21.9´n 29°40.1´w). mountain in paul stern land, so named by eduard wenk during lauge koch’s 1958 expedition for the overhanging snow masses. snehvide 73ø-325 (73°57.1´n 23°30.2´w). mountain 1394 m high in central hudson land. named by heinrich bütler during lauge koch’s 1938–38 expeditions. snehytten – see arentzhytta. snehætten 72ø-406 (72°39.6´n 26°04.3´w; map 4). mountain in nw lyell land, named snöhättan by a.g. nathorst in 1899 because of the small ice caps which form the summit. (snohátten.) snekollen 73ø-553 (73°01.0´n 27°51.5´w). snow-capped mountain east of mercanton gletscher, goodenough land, named by james wordie’s 1929 expedition as snow dome. it was covered by treacherous deep snow covering wide crevasses in august 1929. (sne kollen knolde.) snekuppel 71ø-301 (71°41.3´n 24°36.5´w; map 5). minor snowcapped summit 1480 m high on the west side of schuchert dal. named by enrico kempter during lauge koch’s 1956–58 expeditions. it has also been called derry. snekuppelkløft 71ø-318 (71°39.0´n 24°31.7´w; map 5). ravine draining south from snekuppel. named by e. kempter during lauge koch’s 1956–58 expeditions. snelejedalen 71ø-97 (71°43.5´n 22°15.4´w). valley in north can ning land, so named during the 1931–34 treårsekspeditionen by arne noe-nygaard because the valley was usually filled with snow. snella 72ø (72°46.1´n 22°51.1´w). small island in vega sund, one of the scott keltie øer. used only on nsiu maps (lacmann 1937), and so named because it resembles in shape the trigger (= snelle) of a gun. snemarken 74ø-383 (74°16.2´n 21°12.6´w). ice cap on central clavering ø. the name (= snow field) was suggested by the place name committee in 1951 as a replacement for the lars christen fig. 81. the inner end of smallefjord, looking northwards to ejnar mikkelsen gletscher and stejlgletscher that drain kong wilhelm land. canongletscher is in the foreground. the john haller photograph collection, geus archive. canongletscher ejnar mikk els en g le ts ch er stejlgletscher smallefjord kong wilhelm land 307 sen fonna of nsiu maps, rejected on the grounds that lars chris ten sen was then still alive. snenæs 76ø-31 (76°49.2´n 19°21.4´w; map 4). peninsula on winge kyst in south germania land, so named by the 1906–08 danmarkekspeditionen because it was usually snow-covered. this point has also been called store snenæs to distinguish it from lille snenæs. (snow naze, sne point.) snenæshytten 76ø-197 (76°49.2´n 19°21.22´w). danish hunting hut at snenæs on the south coast of germania land, built by nanok in august 1933. it has more usually been known as store snenæs hytten to distinguish it from lille snenæshytten. a newer hut nearby is known as ny store snenæs hytte. snerta 73ø (73°52.5´n 20°34´w). small river on the north coast of home forland, so named on an nsiu map (1932a). derived from a norwegian dialect word. (snerta river.) sneryggen 74ø-82 (74°49.2´n 20°06.1´w). snow-covered mountain ridge about 1000 m high on se kuhn ø, named schneerücken by karl koldewey’s 1869–70 expedition. snespurvefjeld 79ø-17 (79°43.0´n 20°52.0´w; map 4). mountain near kap bernhoft in south kronprins christian land. named by the 1938–39 mørkefjord expedition for the snow bunting (plectro phenax nivalis). (snespurvefjæld.) snestormdal 73ø-371 (73°53.0´n 26°06.0´w; map 4). valley in north andrée land between eremitdal and nunatakgletscher. named during lauge koch’s 1948–50 expeditions by erdhart fränkl, who experienced an unpleasant snowstorm here early one summer. snesund 70ø-25 (70°49.0´n 27°15.0´w; maps 3, 4). sound between storø and milne land. so named by carl ryder’s 1891–92 expedition because 2–3 feet of loose snow were encountered here during their first winter journey in april 1892. snesø 70ø-375 (70°17.0´n 29°00.0´w; map 4). lake in west gåseland on the north side of vindblæsedal, so named by eduard wenk during lauge koch’s 1958 expedition. snetoppen 71ø-300 (71°57.1´n 25°17.8´w; map 5). mountain 2763 m high between the heads of canta bræ and krabbegletscher, central stauning alper, with a snow summit. named by john haller following explorations during lauge koch’s 1954 expedition, and first climbed by the 1963 cambridge university expedition which reached the summit on skis on 8 august. sneugleelv 74ø-103 (74°08.3´n 20°40.5´w). small river on se clav ering ø, a tributary of fossilelv. named by lauge koch’s 1929–30 expeditions in the form snow owl river, for a sighting of owls. snevigen 71ø-96 (71°43.9´n 22°16.7´w). bay in north canning land, so named during the 1931–34 treårsekspeditionen by arne noe-nygaard for its position at the mouth of snelejedalen. snevigen 71ø (c. 71°44´n 22°16´w). norwegian hunting hut in sne vigen, north canning land, said to have been built in 1932 by helge ingstad’s expedition. no trace of it remains. snevæggen 72ø-396 (72°05.8´n 23°24.1´w). mountain wall in north scoresby land ne of segldal. named by hans kapp during lauge koch’s 1957–58 expeditions (snevæggen = the snow wall). snippen 70ø-234 (70°47.0´n 21°38.6´w). narrow peninsula on the north side of vejle fjord, on the east coast of liverpool land. snow comb 71ø (71°38.2´n 25°19.6´w; map 5). mountain 2272 m high on the ridge between mercurius gletscher and oxford glet scher, south stauning alper. named and first climbed by the 1975 scottish scoresby land expedition led by e.a.m. walker. (comb peak.) snubba 73ø (73°52.5´n 20°28´w). small river on the north coast of home forland, so named on an nsiu map (1932a). snuden [ukaleqartip oqqummut nuaa] 70ø-219 (70°41.0´n 21°34.9´w). peninsula on the east coast of liverpool land. named during the 1931–34 treårsekspeditionen by laurits bruhn with the adjacent features gabet and hagen for the shape on the map (snuden = the snout, nose). snyder bugt 71ø (71°33.0´n 22°51.8´w). name used on 1952 wac maps for the bay on the north side of nordvestfjord at the front of borgbjerg gletscher. origin unknown. snævringen 72ø-331 (72°45.0´n 23°01.0´w). narrow sound be tween kista ø and traill ø, vega sund. the name was proposed by søkortarkivet in 1956–57 following surveying of the channel through vega sund as an alternative approach for ships en route to mestersvig airfield and nyhavn. snøheim 72ø (72°52.7´n 24°01.7´w). norwegian hunting hut on the south side of vega sund, north of rebild. built by arktisk nærings drift in august 1929, it is now a ruin. it has also been known as østhytta and traill hytten (snœheim.) soel-backen 76ø (76°56.1´n 21°28.4w). mountain ridge 690 m high west of danmarks monumentet on the south side of mørkefjord, the present redekammen. the name appears only on the christ mas card sent to peter freuchen at pustersvig in 1907 during the 1906–08 danmark-ekspeditionen, and is a variation of solbakken (= sunny hillside). the card is reproduced in koch (1912, 1916). sofia sund 72ø-65 73ø-274 (73°02.0´n 23°50.8´w; maps 3, 4). sound between ymer ø and geographical society ø. named as sofias sund by a.g. nathorst in 1899 after the ship sofia, which carried swedish expeditions to spitsbergen in 1868 and to greenland in 1883. (sofia strait, sofia sound, sofiasund.) sogneelv 73ø-195 (73°39.9´n 21°39.5´w). river in west hold with hope, flowing into loch fyne near botnhuset. adapted from the original sokna on the 1932a nsiu map. both norwegian and danish words translate as parish or district. solbakken 74ø-388 (74°03.7´n 26°41.6´w). east end of bernhard studer land, so named by hans r. katz during lauge koch’s 1951 expedition because they had a campsite here, a warm and sunny location. soldal 72ø-476 (72°06.1´n 26°20.3´w). south-facing glacier-filled valley north of the nw end of furesø, nathorst land, named by hans zweifel during lauge koch’s 1954–55 expeditions (soldal = sun valley). solfaldsdal 71ø-77 (71°45.9´n 23°00.0´w). valley on the nw side of fleming fjord. so named by arne noe-nygaard during the 1931–34 treårsekspeditionen because the sun set here as seen from vimmelskaftet station. solgletscher 70ø-86 (70°13.1´n 24°30.3´w; map 4). glacier on volquaart boon kyst west of soltemplet, so named during the 1931–34 treårsekspeditionen by laurits bruhn. solheim 73ø (73°32.9´n 24°25.0´w). name used by orvin (1930) for a prospective hunting hut at kap ovibos, se strindberg land. material for the hut was put ashore here by arktisk næringsdrift in 1931. it was probably named after wilhelm solheim. see also solheimfjellet. the present hut on the site is known as kap ovibos hytten. solheimfjellet 73ø (73°50.5´n 20°44.6´w). mountain 1015 m high in home forland. so named on an nsiu map (1932a), after wilhelm solheim [b. 1890], an nsiu surveyor who took part in 20 expeditions to the arctic, including the 1929 and 1931–33 nsiu expeditions to east greenland. solifluktionsgletscher 74ø-311 (74°05.8n 21°14.1w). name used for two areas east of eskimonæs, southern clavering ø, which are characterised by mud flows produced by solifluction. solitaryplateau 74ø-229 (74°07´n 20°46´w). small plateau on se clavering ø, so named during lauge koch’s 1929–30 expeditions because it is isolated by a fault from adjacent parts of the plateau. solitærbugt 72ø-123 (72°52.8´n 25°06.4´w; map 4). bay on north ella ø, on the west shore of which lauge koch’s scientific station and the sirius depot houses are situated (fig. 40). the name is said to have originated from aage de lemos, telegraphist on ella ø during the 1931–34 treårsekspeditionen and refers to the idealistic setting of the station (solitær = diamond). see also kystens per le. use of the name in zoological reports led to its formal adoption. 308 solstrand 73ø-601 (75°34.6´n 24°42.3´w). south-facing beach on the coast of south strindberg land. named for its sheltered setting, delightfully warm on sunny days. the name was first used as a botanical reference locality in reports of the 1931–34 treårseks peditionen (gelting 1934). solkæret 74ø (74°28.2´n 20°35.4´w). reference locality west of zackenberg forskningsstation, used in reports by visiting scientists. solstrand 73ø (73°48.4´n 24°02.2´w). norwegian hunting hut on the east side of waltershausen gletscher, 8 km north of kap bull. it was built in 1938 for ole klokset’s expedition, and has also been known as rødtophytten and brehytta. solstrand 72ø (c. 72°13´n 23°45´w). norwegian hunting hut near noret, built in august 1930 for the møre expedition and originally called lavøira. it was moved in 1954 to fleming fjord. solstrand 75ø (75°32.8´n 21°28.1´w). norwegian hunting hut on the east side of brædal in bredefjord, built for john giæver’s expedition in august 1933. it was also known as brædalhytten. no trace of the hut remained in 1988. solstranda 72ø (72°47.9´n 22°46.5´w). west-facing coastal stretch of geographical society ø, on the east side of central vega sund. so named on the nsiu maps of lacmann (1937) because it has a pleasant sheltered beach, a sun-trap in good weather. soltemplet 70ø-87 (70°12.0´n 24°21.3´w; map 4). mountain on volquaart boon kyst between månegletscher and solgletscher. so named by laurits bruhn during the 1931–34 treårsekspeditionen for its resemblance to a temple (= sun temple). solvefjellet 73ø (73°25.7´n 23°14.1´w). mountain on the south side of gauss halvø, the south end of the present stensiö bjerg. so named on an nsiu map (1932a), after solvi (or solve), one of the original viking settlers of greenland. (mt. solve.) solveigs hytta 72ø (72°51.8´n 23°33.7´w). original name of the nor wegian hunting hut built in august 1929 for arktisk nærings drift on the north side of vega sund. it was named by thor halle after his wife. the hut has also been known as revodden, kapp rygg and sverdrup hytte. solveigs sang 71ø (71°54.2´n 25°07.0´w; map 5). snow summit about 2410 m high on the north side of roslin gletscher, between fimbulbren and valhalbreen. climbed by the 1996 norwegian stauning alper expedition, and named after ‘solveigs sang’ from peer gynt by henrik ibsen. solvig 80ø-123 (80°30.0´n 20°13.2´w; map 4). inner n–s-trending branch of ingolf fjord. named during operation groundhog 1960 together with månevig (sol = sun, måne = moon). sommerfugle sø 74ø (74°29.9´n 20°36.2´w). small lake in the area known as morænebakkerne, north of zackenberg forsknings station. the name is used as a reference locality by scientists studying lake ecosystems. (sommerfuglesø.) sommerterrassen 80ø-55 (80°39.2´n 15°25.3´w). terrace about 4 m above sea level on the se coast of amdrup land, north of kap jungersen. named by the 1938–39 mørkefjord expedition for the presence of an inuit summer settlement of tent-rings and meat caches. sonja havn 76ø-72 (76°36.5´n 18°36.4´w). small harbour on the south coast of lille koldewey found by alf trolle in july 1907. so named during the 1906–08 danmark-ekspeditionen by christian b. thostrup after his daughter (thostrup 2007). (sonjas havn, sonya harbour.) sonklardal 73ø (73°07.1´n 26°05.9´w). name occasionally used by haller (1955) for the valley in ne suess land occupied by sonklar gletscher. sonklargletscher 73ø-510 (73°07.1´n 26°05.9´w; map 4). glacier on the south side of kejser franz joseph fjord, named by karl koldewey’s 1869–70 expedition as sonklar gletscher for karl son klar [1816–1885], lecturer in geography at the military academy in wiener-neustadt (j. løve, personal communication 2010). the glacier was climbed by julius payer in the summer of 1870 to reach the viewpoint from which petermann bjerg was seen for the first time. sonnblick spids 72ø-508 (72°03.8´n 25°15.5´w; map 5). moun tain on the north side of sefström gletscher, first climbed and so named by hans gsellman’s 1957 expedition. they had also called it dritten lagergipfels. the name was approved at the suggestion of the 1963 cambridge university expedition. (sonnblickspitze, sonn blick.) sonnenjoch 75ø (c. 75°19´n 17°50´w). feature in the vicinity of the base camp of the 1943–44 operation bassgeiger at kap sussi, shan non. the name is recorded by olsen (1965). sonnenkopf 74ø-41 (74°40.2´n 18°26.5´w). highest mountain (602 m high) on lille pendulum. so named by karl koldewey’s 1869–70 expedition, possibly for a mountain of similar name in the austrian alps. (mt sonnenkopf.) sophie holm 79ø-18 (79°55.2´n 17°20.6´w). small island off the east coast of hovgaard ø, south of kap h. n. andersen. named during the 1938–39 mørkefjord expedition after hovgaard’s wife, sophie christiane nielsen [1856–1934]. eigil knuth visited the island in june 1939. (sofieholmen, sophies holm.) sophus müller næs 80ø-13 (80°47.1´n 14°08.5´w; maps 1, 4). cape in eastern amdrup land. named by christian b. thostrup as sophus müllers næs during the 1906–08 danmark-ekspeditionen after sophus müller [1846–1934], a danish archaeologist who was director at the national museum in copenhagen. the name is sometimes applied to the more prominent slightly more northern cape. (sophus müllers naze.) soppbukta – see svampebugt. soranerbræen 76ø-18 (76°07.0´n 22°00.0´w; maps 2, 4). glacier draining into the sw part of dove bugt between ad. s. jensen land and rechnitzer land, so named by the 1906–08 danmarkekspeditionen. a ‘soraner’ is a student of sorø akademis skole in denmark. henning bistrup, one of the expedition, studied here from 1890 to 1893. (soranergletscher, soraner glacier, soranerjö kull.) soria moria 73ø (73°54.0´n 24°24.1´w). name proposed by the norwegian hunter john giæver in 1930 for the most distant nuna taks in waltershausen gletscher. it is the name of a castle in a nor wegian fairy-tale. giæver (1931) suggested it bears comparison with the faraway how nunataks in wordie gletscher. sorte hjørne 73ø-380 (73°40.0´n 25°05.8´w). cape on the west side of geologfjord, south of the mouth of morænedal, east andrée land. named during lauge koch’s 1948–50 expeditions by erdhart fränkl for the colour (sorte = black). (kap sorte hjørne.) sorte hjørne – sortehjørnehytten. sorte kløft 75ø (75°09.4´n 19°55.3´w). name used by danish hunters for a minor ravine draining part of søndre muschelbjerg, hochstetter forland (nyholm-poulsen 1985). sorte knold 75ø (75°10.5´n 19°58.3´w). name used by danish hun ters for a minor feature near the coast south of jarners kulmine, probably identical with negeren. sorte odde 72ø (70°42.9´n 27°39.4´w). name used in the diaries of helge vedel (gulløv 1991) during carl ryder’s 1891–92 expedition, apparently for the south point of storø. sorte pynt 70ø (70°31.0´n 28°21.0´w). name used by carl ryder’s 1891–92 expedition for a locality in vestfjord, probably identical with the present kobberpynt. nordenskjöld (1907) in his description of a sample collected here refers to the locality as sorte pynt or black point. sortebakker 80ø-31 (80°10.6´n 17°16.3´w; map 4). coastal mountains west of depotfjeld in south holm land. so named by eigil nielsen during the 1938–39 mørkefjord expedition, because of the occurrence of black coal seams. (sortebakkerne.) sortebjerg 72ø-218 (72°04.9´n 24°08.5´w; map 5). mountain south of the mouth of nedre funddal, north scoresby land. 309 named by prospecting teams associated with lauge koch’s 1948– 49 expeditions. sortehjørne has been used for the mountain (pessl 1962), and is commonly used for the nearby hut. sortebjerg hytte – see sortehjørnehytte. sortebræ 68ø-21 (69°00.0´n 27°18.0´w). large glacier draining south to the blosseville kyst, so named during g.c amdrup’s coastal survey in a small boat in 1900. the northernmost branches of the glacier extend north of latitude 69°n. sorteelv gletscher 72ø-246 (72°20.3´n 24°36.3´w; map 5; fig. 78). glacier in the north stauning alper, named by erdhart fränkl during lauge koch’s 1950–51 expeditions for the colour of the river draining the glacier. sortefjeld 73ø-362 (73°48.8´n 25°17.9´w). mountain in west strind berg land. named during lauge koch’s 1948–49 expeditions by hans r. katz for the colour (sort = black). sortefjeld 77ø-83 (77°19.8´n 21°18.6´w; map 4). mountain north of tværdalen on the west side of annekssøen. named during the 1938–39 mørkefjord expedition as sortefjæld, for its colour, probably by paul gelting who visited it in june 1939. sortefjelde 72ø-236 (72°20.1´n 23°06.4´w). mountain range in se traill ø, so named by desmond t. donovan during lauge koch’s 1949–50 expeditions, for their colour. sortehat 70ø-114 (70°54.7´n 22°48.8´w). mountain west of the head of hurry inlet, named by alfred rosenkrantz during lauge koch’s 1926–27 expeditions as black cap mt, for its appearance. sortehest 72ø-413 (72°21.7´n 28°34.5´w; map 4). isolated nuna taks sw of cecilia nunatak. the name first appears on the 1932 1:1 million scale geodætisk institut map prepared on the basis of aerial observations by lauge koch during the 1931–34 treårseks peditionen, and was given for their appearance (sortehest = black horse). sortehjørnehytte 72ø (72°05.8´n 24°00.8´w). name commonly used for the hut built by nordisk mineselskab in 1952 at the east foot of sortebjerg, sw of mesters vig. it was used by prospecting teams drilling for lead. (sorte hjørne) sorteryg 74ø-215 (73°59.3´n 21°24.4´w). minor ridge in nw hold with hope, beside river 14, covered by black basalt debris. so named by eigil nielsen during the 1931–34 treårsekspeditionen. sorteskær 76ø-162 (76°39.6´n 20°13.6´w). skerry sw of orien terings øerne in dove bugt. discovered and named during the 1932 gefion expedition. (sorte skær.) sorteø 70ø-26 (70°40.3´n 27°43.3´w; map 4). island in rødefjord, named by carl ryder’s 1891–92 expedition as sorte ø, because of its dark colour. (sorte ö, sorte island.) sorthorn 73ø-683 (73°31.1´n 27°20.2´w). mountain 1328 m high in se louise boyd land, west of the front of gerard de geer gletscher. named by john haller following explorations during lauge koch’s 1949–51 expeditions, for the colour of the rocks. south cirque glacier 73ø (73°32.0´n 27°28.9´w). name used in a report by odell (1937a) for the south tributary of louise gletscher, louise a. boyd land, studied during louise boyd’s 1933 expedition. south lochan 72ø (72°14.4´n 23°55.0´w). name used by the univer sity of dundee expeditions between 1968 and 1974 for a small pool near langdyssen at the ne end of mestersvig airfield. southcape 74ø (74°05.3´n 21°17.1´w). name occasionally used in reports of the 1906–08 danmark-ekspeditionen for the south point of clavering ø, the present eskimonæs (thostrup 1911). southern bræ ø 76ø (76°43.5´n 22°06.7´w). most southerly of the bræ øer in borgfjorden. the island was used by the 1952–54 bri tish north greenland expedition as one of their main survey sta tions, and the name appears on the maps of hamilton et al. (1956). southern fault valley – see fault valley and forkastningsdalen. southern ridge – see sønderås. spalenbjerg 73ø-706 (73°11.4´n 29°06.0´w). mountain in south martin knudsen nunatakker, on the west side of victor madsen gletscher. named by john haller and eduard wenk following explorations during lauge koch’s 1951 expedition, after the street spalenberg in the old town centre of basel, switzerland. spaltegletscher 73ø-382 (73°40.4´n 25°18.1´w; map 4). glacier in east andrée land draining via morænedal to geologfjord. named by erdhart fränkl during lauge koch’s 1948–50 expeditions for its crevasses (= spalte). (spalte gletscher.) spaltegletscher 73ø-710 (73°10.1´n 28°27.9´w). tributary glacier to gregory gletscher on its nw side. named during lauge koch’s 1951 expedition by john haller and eduard wenk for the spectacular and very wide crevasses. (spaltengletscher.) spaltegletscher 79ø-20 (79°43.5´n 20°16.0´w; maps 1, 4). branch of the floating e–w-trending glacier filling nioghalvfjerdsfjorden which extends north into dijmphna sund. named by elmar dra strup’s 1938–39 expedition for the numerous wide crevasses. gustav thostrup and alfred wegener were delayed by the crevasses when making the first traverse of the glacier in 1907. spartansletta 72ø (72°42.4´n 22°37.1´w; fig. 14). low lying area on south geographical society ø north of kap hovgaard. used on the nsiu maps of lacmann (1937), the name was given for the spartan aeroplane used for the 1932 nsiu aerial photography. spath fjaeld 70ø (70°38.4´n 22°43.1´w). summit on the west side of hurry inlet between moskusoksekløft and astartekløft. so named by hermann aldinger during the 1931–34 treårsekspeditionen, for l.f. spath [1882–1957]. see also spath plateau. spath plateau 73ø-42 (73°53.8´n 21°27.8´w; map 4). plateau up to 1510 m high in north hold with hope. so named by lauge koch’s 1929–30 expeditions in honour of leonard frank spath [1882– 1957], an english palaeontologist and stratigrapher at the british museum (natural history), who identified many of koch’s fossil collections from the region. the name was apparently originally given to frebold bjerg, the plateau extending from slightly south of kap stosch along the coast eastwards to blåelv (koch 1931), but is now applied to a more extensive plateau 1500 m high and slightly farther south. (spaths plateau.) spejderhatten 73ø-669 (73°43.2´n 27°00.5´w; map 4). mountain in andrée land with the charactersistic shape of a scout’s (= spejder) hat. named by john haller following explorations during lauge koch’s 1949–51 expeditions. spenna 72ø (72°38.7´n 22°24.6´w). island in the east part of vega sund east of nordenskiöld ø. used only on nsiu maps (lacmann 1937), and named after an island of the same name in the troms district of norway. sphinx 70ø (70°46.4´n 26°11.1´w). peak 1920 m high on the south side of korridoren, milne land. climbed by the 2004 west lan cashire scouts expedition. sphinx gletscher 73ø (73°23.6´n 26°21.6´w). glacier draining northwards from the mountain sfinksen in southern andrée land. the name is found on a sketch drawn by john haller in 1949, and published in schwarzenbach (1993). spiralkløft 73ø-564 (73°31.6´n 24°50.8´w). ravine in east andrée land draining via tillitekløft into geologfjord. named by christian poulsen during lauge koch’s 1929 expedition as spiral creek. the position is shown incorrectly on official geodætisk in stitut maps. spiret 70ø-85 (70°15.0´n 24°49.8´w; map 4). mountain on vol quaart boon kyst west of solgletscher, so named during the 1931–34 treårsekspeditionen by laurits bruhn for its appearance (spiret = the spire). spiret 72ø-494 72°07.7´n 24°47.3´w; map 5). dramatic rock peak about 2000 m high at the head of bersærkerbræ, north stauning alper. first climbed by john hunt’s 1960 expedition and named originally in the form bersaerkerspire or bersaerker spire, names still commonly found in mountaineering literature in preference to the official name. splinten 77ø-46 (77°14.4´n 24°27.8´w; fig. 21). prominent rock 310 ridge west of prins axel nunatak, dronning louise land. named by the 1909–12 alabama expedition for the shape (splinten = the splinter). the ridge was traversed by members of the 1952–54 british north greenland expedition. splinten col 76ø (77°11.5´n 24°30.0´w). this name was occasionally used by members of the 1952–54 british north greenland expedition for the col south of splinten, dronning louise land (simp son 1957). splitbæk 81ø (81°14.9´n 13°36.0´w). stream in nw kilen, kron prins christian land, with a fanning and anastamosing course. the name is found on a coloured geological map of kilen printed in 1991 (pedersen 1991). sporfjeld 71ø-170 (71°53.3´n 22°46.8´w). mountain north of the mouth of ørsted dal. so named during lauge koch’s 1938–38 expeditions by hans stauber, because he had found geological evidence here for a theory previously doubted (spor = track, clue). sporfjeld hytte 71ø (71°52.2´n 22°45.6´w). hut at the entrance of fleming fjord below sporfjeld. see also lapstun-hytten. sporvognen 72ø (72°28.6´n 24°01.5´w). hut at the mouth of karupelv, ne of the haslum øer. it was moved to this site from mestersvig airfield in 1976, and was intended as a holiday hut for airfield personnel. reported in good condition in 1990. spuedalen 74ø (c.74°16´n 19°28´w). name used by the 1908–09 floren expedition, probably for one of the valleys west of kap borlase warren. exact position uncertain. derived probably from the norwegian dialect word for a bird of the curlew family. spurvebugt 70ø-411 (70°31.9´n 26°01.1´w). small bay on sw milne land sw of mudderbugt. named during the 1967–72 ggu scoresby sund expeditions by svend funder for the numerous snow buntings. spydodden 76ø-26 (76°48.6´n 20°46.8´w; map 4). elongate peninsula south of the mouth of hellefjord, so named by the 1906–08 danmark-ekspeditionen because it resembles in shape the point of a spear. a nanok hut a few kilometres to the south at the mouth of port arthur is sometimes known by the name spydodden. (spyd odde, spyde point, spjótsoddi.) spydøen 76ø-288 (76°48.9´n 20°43.6´w). small island north of the north point of spydodden, off east daniel bruun land. named by the 1938–39 mørkefjord expedition. (spydø.) spærrebugt 70ø-210 (70°34.8´n 21°40.5´w). bay on the south side of lillefjord, on the coast of se liverpool land. named by helge g. backlund who explored this region during the 1931–34 treårs ekspeditionen (spærre = obstruct, block). (spoerrebugt.) spærregletscher 71ø-146 72ø-94 (72°00.1´n 25°39.5´w; map 5). large glacier which forms a complete barrier (= spærre) across the east end of furesø. named by ove simonsen during the 1931–34 treårsekspeditionen. st. andrews klippe 76ø-330 (76°33.8´n 25°36.6´w; map 4). large cliff on the north side of budolfi isstrøm, dronning louise land. named by the 1952–54 british north greenland expedition after the university of st andrews, scotland’s oldest university founded in 1410 as st. mary’s college. one of the expedition members, peter wyllie, had graduated from here. recent official names lists have deleted the ‘s’ in ‘andrews’ (st. andrew klippe.) st. bartholomews tårn 72ø-511 (72°03.1´n 24°56.1´w; map 5). rock peak with twin summits about 2440 m high sw of crescent pas, stauning alper. named by the 1963 cambridge university expedition who climbed it on 23 august. (st. bartholomew’s tower, torre di s. bartolomeo.) st. johns tinde 72ø-504 (72°05.5´n 25°08.2´w). peak 2200 m high on the ne side of cavendish gletscher, stauning alper. climbed by the 1963 cambridge expedition on 20 august, and named after st john’s college, cambridge, founded in 1511 on the site of the hospital of st. john. official name lists omit the genetive ‘s’. st. petersburg bjerg – see mount petersberg. stabbene 74ø (c.74°16´n 19°23´w). name used by the 1908–09 floren expedition for basalt columns in the vicinity of kap borlase warren. the norwegian word translates as something short or stub by. exact position uncertain. stakkarsdalen 74ø (74°25.5´n 19°22.4´w). this name was apparently used by the 1908–09 floren expedition for dronning augusta dalen in wollaston forland (brandal 1930). stakkeløbet 76ø-278 (76°29.6´n 20°44.2´w). sound between god fred hansen ø and stakken in the sw part of dove bugt. named by the 1938–39 mørkefjord expedition (stakke = haystack). stakken 74ø (c. 74°10´n 20°12´w). name used by 1927–29 hird expedition for a feature in the vicinity of the hunting station at kap mary, east clavering ø (rogne 1981). stakken 76ø-277 (76°30.0´n 20°40.9´w; map 4). small island east of godfred hansen ø. named by the 1938–39 mørkefjord expedition for its appearance (stak = haystack). startdal 72ø-146 (72°13.1´n 22°26.8´w). valley on se traill ø east of the head of drømmebugten. so named during lauge koch’s 1938–38 expeditions by hans p. schaub, probably because he started his geological work here. station ‘a’ – see carlshavn. station ‘b’ – see kap broer ruys station. station-mountain 69ø (69°24.7´n 24°04.0´w). the name was used only by böggild (1905), for the 1300 m high summit of kap dalton. see also stationsbugt. stationsbugt 69ø (69°26.0´n 29°07.0´w). small bay north of kap dalton. the antarctic anchored here in july 1900 during g.c. amdrup’s 1898–1900 expedition while a depot house (amdrup hytte), and possible wintering station, was built on land. this name is only used by jacobsen (1900). stationsø 76ø-182 (76°02.4´n 19°57.8´w). small island sw of kap beurmann at the mouth of bessel fjord. it was used by thostrup (1911) as a reference locality in his archaeological report of the 1906–08 danmark-ekspeditionen. statuebjerg 70ø-109 (70°50.4´n 22°50.4´w). mountain on the west side of the head of hurry inlet, named by alfred rosenkrantz during lauge koch’s 1926–27 expeditions as statue mt for its appearance. stauning alper 71ø-144 72ø-80 (72°00.0´n 25°00.0´w; maps 3, 5; figs 27, 31, 38). alpine mountain range bounded to the north by kong oscar fjord and segelsällskapets fjord, to the east by skel dal, schuchert flod and holger danske briller, to the west by alpefjord and borgbjerg gletscher, and to the south by part of nordvestfjord. the range was observed and partly mapped by early explorers (see rink bjerge), but first seen completely from the air by lauge koch in 1932. it was named after thorvald august marinus stauning [1873–1942], noted danish politician and prime minister for 15 years, who, koch reports, always took great interest in his work in east greenland and was ever ready to grant financial aid. the region has been extensively explored by climbing expeditions. (stauning alps, stauningalpen.) steensby bjerg 73ø-51 (73°54.9´n 21°04.8´w). mountain in north hold with hope, named by lauge koch’s 1929–30 expeditions in the form mt steensby for hans peder steensby [1875–1920]. steens by was a danish geographer and professor at the university of copenhagen, with interests in north africa and the ethnology of polar inuit. steenstrup bjerg 72ø-58 (72°17.7´n 22°51.9´w). mountain 1294 m high on se traill ø. named steenstrups berg by a.g. nathorst in 1899 after johannes japetus smith steenstrup [1813–1897], with whom nathorst cooperated on studies of glacial flora as a student. japetus steenstrup was a noted danish natural historian and professor of zoology at the university of copenhagen from 1846 to 1885, who made significant research in zoology, botany and archaeology. (steenstrup mountain, mt. steenstrup.) steenstrup dal 72ø-148 (72°16.2´n 22°56.8´w). valley in extreme se traill ø, south of steenstrup bjerg. named during lauge koch’s 311 1938–38 expeditions by hans p. schaub. see also steenstrup bjerg. steffensens hytte 73ø (73°10.6´n 23°08.3´w). norwegian hunting hut at the mouth of dusén fjord, built by arktisk næringsdrift in 1929. it has been more commonly known by the names kikut and dusens fjordhytten. (steffensen.) stegocephalryg 74ø-226 (74°01.6´n 21°36.3´w). minor ridge north of frebold bjerg, adjacent to river 7, nw hold with hope. so named by eigil nielsen during the 1931–34 treårsekspedi tionen, for the find of a particularly fine example of the fossil fish ‘stegalocephalus’. steinbjerg 71ø (71°47.9´n 24°58.0´w; map 5). mountain about 1950 m high on the sw side of roslin gletscher. climbed by karl herlig koffer’s expedition on 21 august 1966. steinenbjerg 73ø (c. 73°10´n 29°05´w). name used by buess (1953 p. 216) for a mountain in the martin knudsen nunatakker. it was named during explorations on lauge koch’s 1951 expedition after the street steinenberg in the old town centre of basel, switzerland. steinmannspids 74ø-76 (74°10.7´n 20°49.3´w). mountain 1332 m high on clavering ø. named steinmann spitze by karl koldewey’s 1869–70 expedition, because of the cairn-like basalt pillars which crown the summit. (steinmannen, mt. steinmann, steinmann peak.) steinröysa 73ø (73°11.5´n 22°56.0´w). small island east of the vinterøer, at the mouth of dusén fjord. so named on an nsiu map (1932a). the name implies a stony desert or plain. steinrøisdal 71ø (71°55.8´n 23°58.6´w). valley south of antarctic havn in scoresby land, the present flexurdal. the name has been often used by norwegian hunters (ingstad 1937), and appears on norsk søkort 511 (1937). moskusdal has been used for the same valley. steinrøysdalshytta 71ø (71°53.1´n 23°01.0´w). name sometimes used for the norwegian hunting hut built by helge ingstad’s 1932–34 expedition in henrik møller dal, close to the junction with flex urdal (steinrøisdal). it is also known as minimalen and øyedals hytten. steinsund 73ø (73°58.8´n 21°08.9´w). narrow sound between strip öya and vesle finschöya in the finsch øer group. so named on an nsiu map (1932a). stejlfjeld 70ø-84 (70°17.8´n 24°44.2´w; map 4). steep cliff on volquaart boon kyst west of solgletscher. so named by laurits bruhn during the 1931–34 treårsekspeditionen for its appearance (stejl = steep). engelsborg has also been used. stejlgletscher 73ø-635 (73°05.0´n 26°45.5´w; map 4). small glacier in nw suess land. so named by ove simonsen during the 1931–34 treårsekspeditionen because it descends steeply to kejser franz joseph fjord (stejl = steep). stejlgletscher 74ø-374 (74°40.0´n 22°09.9´w). glacier in west a.p. olsen land on the ne side of tyrolerdal. so named by the 1948 leeds university expedition because it was one of the steepest on which they worked. (steep glacier.) stejlgletscher 75ø-74 (75°34.0´n 22°38.8´w; fig. 81). glacier west of the head of smallefjord. the name originated from the wintering party at kulhus during the 1931–34 treårseks peditionen. stejlpynt 70ø-428 (70°26.1´n 26°56.3´w). cliff on the south side of fønfjord with a near vertical face. named during the 1967–72 ggu scoresby sund expeditions by georg sawatzki. stendal 71ø-434 (71°08.1´n 28°47.5´w). valley in graben land, characterised by abundant loose boulders. named by peter home wood during the 1967–72 ggu scoresby sund expeditions. stendysse bjerg 74ø (c. 73°53´n 20°56´w). name used in a report on 1938 field work during lauge koch’s expedition (in: koch 1955 pp. 586, 589) for a mountain east of kap stosch adjacent to diener bjerg (stendysse = cairn, burial mound). stenen 78ø-29 (78°38.2´n 23°05.2´w; map 4). northern peak of moltke nunatak. the name was given by the place name com mittee in 1940 to replace a suggestion by the 1938–39 mørkefjord expedition. an association with milepælen immediately to the south was intended (stenen = the stone). stenknolden 76ø (c. 76°49´n 18°19´w). marked feature on the coast of east germania land south of syttenkilometernæs; exact position uncertain. the name was used by thostrup (2007) in his account of the 1906–08 danmark-ekspeditionen (j. løve, personal communication 2009). stenløsgletscher 73ø-613 (73°12.4´n 28°03.8´w). glacier on the south side of knækdalen, notable for the absence of moraine or dirt bands. named by louise boyd in 1933 as moraineless glacier. stenmanden 73ø-678 (73°35.1´n 26°22.9´w). mountain at the west end of grejsdal, where it divides into gnejsdal and djævlekløften. named by john haller following explorations during lauge koch’s 1949–51 expeditions, for the presence of a large band of migmati tic gneiss in the shape of a man. steno bræ 69ø-23 (69°51.0´n 23°40.3´w; maps 3, 4). glacier west of manby halvø on the north blosseville kyst. named by g.c. amdrup’s 1898–1900 expedition after the pioneer danish geologist niels steensen. nicolaus stenonis (niels steensen) [1638–86], a danish cleric, physician and geologist, was noted especially for his ‘prodrome’, an early landmark in the history of stratigraphy, structural geology and palaeontology. (stenos bræ.) steno land 74ø-137 (74°16.3´n 23°50.1´w; maps 2, 4). land area between vibeke gletscher and wordie gletscher, so named by lauge koch’s 1929–30 expeditions. in its original usage the name covered the area extending west to waltershausen gletscher and thus included the present ole rømer land. the limits were more precisely defined as a result of lauge koch’s aerial observations in 1932 (fig. 15). see also steno bræ. stenpikkerelv 71ø-424 71°00.4´n 27°51.5´w). river on the west side of rypefjord. named during the 1967–72 ggu scoresby sund expeditions by svend funder for the numerous wheatear (= stenpikker, i.e. oenanthe oenanthe). stensiö bjerg 73ø-111 (73°25.7´n 23°14.1´w). mountain on the sw coast of gauss halvø. named during the 1931–34 treårseks peditionen by gunnar säve-söderbergh as mt. stensiö, for erik a:son stensiö [1891–1984]. an eminent palaeozoologist, he be came professor at the swedish museum of natural history at stockholm. norwegian maps of the 1930s used solvefjeld for the same feature. stensiö plateau 73ø-50 (73°57.8´n 21°20.0´w). plateau in north hold with hope between gulelv and blåelv, named during lauge koch’s 1929–30 expeditions after erik a:son stensiö who de scribed the fossil fishes collected from the region by koch’s expeditions. (stensiø plateau, stensiöfjellet, stensiöberg.) stensund 71ø-117 (71°19.7´n 21°47.5´w; map 4). fjord in east liverpool land nw of kap topham. so named by helge g. back lund during the 1931–34 treårsekspeditionen, who thought it to be a sound. stenørken 80ø-81 (80°04.3´n 20°33.4´w; fig. 24). plateau in south kronprins christian land between rivieradal and sæfaxi elv. named during lauge koch’s 1952–53 expeditions by erdhart fränkl (stenørken = stony desert). (stenørkenen.) steward ø [sulussuutikajik] 69ø-4 (69°54.3´n 22°52.0´w; maps 3, 4). small island sw of kap brewster. named by william scores by jr. in 1822 as steward island, after charles steward of yarmouth, a companion on one of his earlier voyages to the whale fishery. the name appeared on the maps of the 1879 ingolf expedition (mourier 1880) in error as stewart ø, and subsequently on many other maps in the same form, possibly due to confusion with kap stewart (which has also been misspelt ‘steward’). the german edition of scoresby’s narrative uses the ‘stewart’ form for the cape and island in his appendix (scoresby 1825 p. 414), and ‘steward’ for both features on the chart. stewart ø is commonly used today by danes at scoresbysund. a house was built in a bay on the south side of the island for bear hunting in 1971 on the initiative of jakob sanimuinaq, and a second house added in 1972 (see tsulitsuuligai). 312 stigbøjlen 78ø-25 (78°13.4´n 19°04.2´w; maps 1, 4). large island in jøkelbugten. named by the 1938–39 mørkefjord expedition, together with the adjacent islands hammeren and ambolten, for a supposed resemblance in shape to bones in the ear (stigbøjlen = stirrup). stigdalen 72ø (72°55.6´n 24°08.5´w). valley on west geographical society ø, draining south into vega sund. so named on the nsiu maps of lacmann (1937) because the valley is steep (= stig). stille ø 73ø-45 (73°57.9´n 21°10.3´w; map 4). southern island of the finsch øer group. named by lauge koch’s 1929–30 expeditions in the form stille island for the german petrographer and structural geologist hans stille [1876–1966], noted for his studies of mountain building processes. norwegian maps have used stille öyane to include this and the adjacent small islands, and kilöya for the present stille ø. stirling fjeld 72ø-490 (72°09.8´n 24°31.1´w; map 5). mountain 1640 m high on the south side of bersærkerbræ, north stauning alper. first climbed by john hunt’s 1960 expedition, and at the request of the boys of the expedition named stirling for stirling castle, an outstanding example of renaissance achitecture dating mainly from the 15th and 16th centuries. the second ascent was by the 1963 imperial college expedition. stjernefjeldene 76ø-50 (76°54.5´n 20°53.6´w). range of hills in daniel bruun land south of mørkefjord, named by the 1906–08 danmark-ekspeditionen as stjerne fjeldene. when all the watches at the meteorological station manned by peter freuchen at pusters vig became unreliable, alfred wegener and j.p. koch made a local calculation of star-time, based on the disappearance of a certain star behind this mountain as seen from the door of freuchen’s station in pustervig (koch 1912). (montes stellarum, stjærne fjældene, sternenwände, stjerne mts.) stjernesøen 76ø-298 (76°56.3´n 20°15.5´w). star-shaped lake between lakseelven and mørkefjord station. named by the 1938–39 mørkefjord expedition. (stjærnesøen.) stopklodsen 79ø-21 (79°52.4´n 20°09.1´w). small island in dijm phna sund at the front of spaltegletscher. the name is attributed to david malmquist following his work with lauge’s koch geological expeditions, and arose because the glacier front rests on the island (stopklodsen = door stop). it was approved in 1958. stor-dalen 73ø (73°04.6´n 23°43.1´w). name sometimes used for the norwegian hunting hut east of the mouth of barnabas dal, south ymer ø, and also used by norwegians for the valley itself. the hut was moved to this site from the opposite side of the fjord in october 1930, where it was known as bødtkers hytta, and in 1931 again moved to renbugten. a later hut (originally strømhytten), was moved in august 1932 to this site, where it became known as raudalshytta or barnabasdal hytte. (stordalen, dal hytten.) storborgen 75ø-73 (75°32.3´n 22°00.0´w; map 4; fig. 51). penin sula between smallefjord and bredefjord. the name is attributed to the wintering party at kulhus in 1935, and derives from its resemblance to a large castle (borg = castle). storbræ 68ø-20 (69°00.0n 26°07.0w). large glacier on the blosseville kyst that extends northwards beyond latitude 69°n. storbukta 74ø (74°06.8´n 20°53.6´w). alternative name for døde mands bugten on se clavering ø, occasionally used by norwegian hunters. stordal 73ø-38 (73°45.0´n 22°23.8´w; map 4). large valley in east hudson land draining south into the head of moskusoksefjord, named by james wordie in 1926 as great valley for its size. it has also been called granite valley. stordalen 73ø (73°45.8´n 24°48.8´w). name used by arne høygaard and martin mehren in 1931, and sigurd skaun and harald welde in 1932, for brogetdalen on strindberg land. it has also been called giæverdalen by norwegians. stordalen – see stor-dalen. stordalen 74ø (74°24.2´n 19°09.5´w). name used for the norwegian hunting hut built in july 1928 by the hird expedition at the mouth of dronning augusta dal, wollaston forland. norwegians also used the name stordalen for the valley. it has also been known as augustadalhytten and bjørnebu. stordalshytten 74ø (74°38.5´n 20°49.5´w). norwegian hunting hut on the south side of lindeman fjord, built in august 1932 for sigurd tolløfsens expedition by johan stordal. it is also known as svendsby. store blydal 72ø-193 (72°11.9´n 24°06.3´w; map 5). valley in north scoresby land draining north. named by prospecting teams associated with lauge koch’s 1948–49 expeditions for the major finds of lead ore at blyklippen on the west side of the valley. store bælt 76ø-75 (76°20.0´n 19°30.0´w; maps 2, 4). broad sound west of store koldewey. the name was used in the hydrographical reports of the the 1906–08 danmark-ekspeditionen expedition in the form store-bælt, and was given for the sound of the same name between fynen and sjælland in denmark. see also lille bælt. (storebælt, store belt.) store finsch 74ø-90 (74°02.5´n 20°53.5´w; map 4). largest island of the finsch øer group, first distinguished from the other islands as great finsch island by james wordie in 1926. (store finschöya.) store koldewey 75ø-87 76ø-38b (76°15.0´n 18°42.0´w; maps 2, 4). largest of the koldewey øer. karl koldewey’s 1869–70 expedition refers to an island as grosse koldewey-insel in the astronomy section of the narrative, but may not have intended it as a formal name. the present island was shown on koldewey’s maps as three islands, which the 1906–08 danmark-ekspeditionen showed to be connected and gave the present name to the long narrow island. (great koldewey island.) store myteklippe 70ø-378 (70°14.7´n 29°00.4´w). cliff on the south coast of kaskadesø, west gåseland. so named during lauge koch’s 1958 expedition by eduard wenk, because it and the adjacent cliff (lille myteklippe) were similar in their form and tectonic relationships to the grossen mythen and kleinen mythen in canton schwyz, switzerland. store norske ø 79ø (79°05.0´n 17°48.8´w). name occasionally used in the accounts of the 1938–39 mørkefjord expedition for the largest of the norske øer. store raset 73ø (73°24.6´n 23°15.0´w). fossil locality on the south slope of stensiö bjerg, gauss halvø. the name was used by gunnar säve-söderbergh during the 1931–34 treårsekspeditionen. store ravnefjeld 71ø (71°42.3´n 22°41.5´w). name used by gras mück & trümpy (1969) for the main peak of ravnefjeld on their map of wegener halvø. store sneleje 74ø-312 (74°05.9´n 21°16.6´w). depression north of eskimonæs station, the site of a small stream, often snow-filled. the name originated from the wintering party at eskimonæs during the 1931–34 treårsekspeditionen. store snenæshytten 76ø (76°49.2´n 19°21.2´w). danish hunting hut at snenæs on the south coast of germania land, built by nanok in august 1933. it has also been known as snenæshytten. a new hut was built on the same site in 1999 by danmarkshavn weather station. (store snenæs hytten.) store sø 77ø (77°04.5´n 20°50.4´w). original name for the present sælsøen, which was discovered during the 1906–08 danmarkekspeditionen. this version of the name is mentioned by trolle (1909), and also appears on a sketch map by c.s. poulsen published in lundbye (1984). laksesø has also been used. store sødal 74ø-199 (74°31´n 21°00´w; map 4). valley containing two large lakes, situated north of and parallel to tyrolerfjord. the name first appeared in a botanical report by gelting (1934) on work during the 1931–34 treårsekspeditionen as great lake val ley. (store södal.) store vinteröya 73ø (73°13.0´n 23°07.0´w). larger of the two vinter øer at the mouth of dusén fjord. so named on the 1932a 313 nsiu map for its relative size. storedal 72ø (72°05.3´n 23°58.0´w). term employed by pessl (1962) for the present deltadal, which drains into mesters vig. storeelv 72ø-127 (72°52.4´n 25°06.9´w). river on nw ella ø, draining into solitærbugt. it was so named by the ella ø wintering party during the 1931–34 treårsekspeditionen, because it is the largest river on the island. storefjord [kangertivit anginersaat] 71ø-129 (71°05.4´n 21°54.6´w; map 4). major e–w-trending fjord which cuts through the mountain range of liverpool land. the name first appeared on a map compiled by janus sørensen in 1928, although the mouth of the fjord had been seen by william scoresby jr. in 1822 (his masclet bay). the fjord was first fully explored by helge g. backlund’s party during the 1931–34 treårsexspedition. storehamrene 70ø-399 (70°55.6´n 27°22.1´w). mountain massif in sw renland, on the north side of øfjord. named during the 1963 geodætisk institut expedition after the cliffs of the same name in bornholm. storelv 73ø-329 74ø-336 (73°45.0´n 22°23.8´w). river draining stordal, east hudson land. named by heinrich bütler during lauge koch’s 1936–38 expeditions, and first used in the form stortalfluss. storelvhytten 73ø (73°41.6´n 22°12.7´w). danish hunting hut built in may 1947 for nanok on the north side of storelv. it is also known as arvehytten and vuachehytten. storgletscher 71ø-155a (71°57.0´n 24°43.0´w; maps 4, 5). large glacier in the stauning alper draining east to schuchert dal. bjorn jorsalfarers gletscher and langgletscher have been used for the same feature (both as officially approved names), but in 1971 stor glet scher became the only recognised name. storgaard elv 70ø-283 (70°27.6´n 22°37.1´w). small river in se jameson land nw of kap stewart. named by alfred rosenkrantz during lauge koch’s 1926–27 expeditions after einar storgaard [1890–1981]. an eminent danish geographer, storgaard travelled widely in asia and africa and had particular interests in iceland and greenland. (storgaard river.) storholts hus 74ø (74°12.1´n 21°53.5´w). norwegian hunting hut built in august 1954 by sverre storholt for arktisk næringsdrift. it is situated about 5 km se of kap øtker, and replaced an older hut known as nesodden. it is also known as kap øtker hytten. storlandet 77ø-106 (77°19.5´n 21°20.0´w). name given by the 1938–39 mørkefjord expedition for the highland area west of valde mars muren, covering the areas of both present sønder marken and okselandet. storm p. elv 72ø-219 (72°06.0´n 24°00.0´w; maps 4, 5). river formed by the confluence of peter elv and ping elv, sw of mesters vig. named by prospecting teams associated with lauge koch’s 1948–49 expeditions. robert storm petersen [1882–1949], a danish artist and writer always known as ‘storm p.’, was especially noted for his cartoons. (storm p’s elv.) stormbu 73ø (73°28.1´n 21°56.9´w). norwegian hunting hut built by oscar bang and eiliv herdal for arktisk næringsdrift in october 1938, on vestersletten near fløelv. so named because it sur vived a severe storm on the night it was built. bang (1944) reports that two earlier huts on the same site had been blown down. stormbugt 76ø-9 (76°46´n 19°00´w; map 4). bay north of the koldewey øer, nw of kap bismarck. named sturmbai by karl koldewey’s 1869–70 expedition, because the sledge party was de layed by a three-day storm here in april 1870. (stormbugten, storm bay, stormvik.) stormbugthytten 76ø (76°48.9´n 18°59.8´w). danish hunting hut east of the mouth of stormelv, in stormbugt on the south coast of germania land. built in september 1939 for nanok, it has also been known as stormelvshytten and stormely-hytten. stormdal 72ø-161 (72°25.4´n 22°10.5´w). minor valley on east traill ø on the north side of mountnorris fjord. named during lauge koch’s 1938–38 expeditions by hans p. schaub, presumably for a storm experienced while working here. stormdalen 73ø-161 (73°29.5´n 20°46.9´w; map 4). valley in south hold with hope, so named on an nsiu map (1932a; fig. 13). possibly named after erik storm [1904–36], a norwegian pilot. see also stormfjellet. stormdalen 76ø-267 (76°50.1´n 19°01.1´w). valley in south ger mania land, in which stormelv flows. named by the 1938–39 mørke fjord expedition. stormelv 76ø-66 (76°50.1´n 19°01.1´w). river in south germania land draining into stormbugt. so named by the 1906–08 dan mark-ekspeditionen for its proximity to stormbugt. (stormkapelv, storm river, stormelven, stormá.) stormelvshytten, stormely-hytten – see stormbugthytten. stormfjellet 74ø (74°23.7´n 20°42.7´w). mountain 1100 m high on north clavering ø. used only on nsiu maps (lacmann 1937), the name was given for erik storm [1904–36], a norwegian pilot who led and organised the 1932 nsiu aerial photography. stormgletscher 75ø-72 (75°40.8´n 22°49.0´w; map 4). glacier west of the head of bredefjord. the name originated from the wintering party at kulhus during the 1931–34 treårseks pedi tionen. stormheimen 75ø (75°03.0´n 17°20.5´w). norwegian hunting hut built in march 1953 about 12 km south of kap pansch, on the east coast of shannon. just after they had built the hut, the three hunters involved experienced a violent snow-storm. stormkap 76ø-65 (76°48.5´n 19°01.7´w). cape on the west side of the mouth of stormelv. given by the 1906–08 danmark-ekspedi tionen because, like karl koldewey, they experienced storms while crossing stormbugt. a.l.v. manniche was stranded here by a severe storm in may 1908. (storm cape, storm kap.) stormlandet 77ø-103 (77°40´n 19°30´w; maps 1, 2, 4). land area between orléans sund and penthievre fjord. so named by the 1938–39 mørkefjord expedition because they experienced repeated heavy gales here, which produced smooth snow-free ice. stormnæs 76ø-64 (76°48.3´n 19°09.9´w; map 4). peninsula in south germania land on the north side of stormbugt, so named by the 1906–08 danmark-ekspeditionen for its proximity to storm bugt. (storm naze, store stormnæs, storm point, stormhöfði.) stormpynt 71ø-35 (71°26.7´n 25°26.8´w). small peninsula on the north side of outer nordvestfjord. named in this form by carl ryder’s 1891–92 expedition, because the expedition sheltered here in a storm during their return from the first exploration of nord vestfjord. stormryggen 71ø-315 (71°59.9´n 23°27.1´w). low-lying dolomite ridge in north scoresby land on the north side of kolledalen. it was near one of hans kapp’s camp sites during lauge koch’s 1957– 58 expeditions (kapp 1960), and presumably named for stormy weather. stormsø 76ø (76°49.1´n 22°27.1´w). lake west of stormelv on the south coast of germania land. the name was used in charles poulsen’s (1991) account of the 1906-08 danmark-ekspedition (j. løve, personal communication 2009). storskærene 77ø-74 (77°31.7´n 19°41.0´w). group of large skerries in skærfjorden, so named during the 1931–34 treårsekspedi tionen by david malmquist. storsletten 74ø-351 (74°34.0´n 19°58.0´w). extensive plain in north wollaston forland, sw of albrecht bugt. named by wolf maync and andreas vischer during lauge koch’s 1936–38 expeditions. storstrømmen 74ø-112 (74°17.1´n 20°28.4´w). river on ne clav ering ø draining into young sund. the name first appears on a sketch map in gustav thostrup’s 1921 logbook in the form stor strøm, and was reported as used by danish hunters. it may originally have been applied to the river occupying grønnedal, south of the present location. 314 storstrømmen 76ø-108 77ø-38a (77°05.0´n 22°30.0´w; maps 2, 4; fig. 21). large glacier flowing south between dronning louise land and daniel bruun land, that merges with l. bistrup bræ to form bredebræ. named by the 1906–08 danmark-ekspeditionen for its size (= the great stream). (storströmmen, storstrommen, stór istraumur.) stortoppen 72ø (72°08.1´n 25°03.3´w). one of the original names used by the norwegian climbers who made the first ascent of nor ske tinden in 1954, the second highest mountain in the stauning alper (hoff 1955). see also eirik raudes tinde. storø [kaasarip nasaa] 70ø-6 (70°49.5´n 27°30.0´w; maps 3, 4). largest of the islands on the east side of rødefjord. discovered and so named by carl ryder’s 1891–92 expedition during their first winter sledge journey. storøen 78ø-21 (78°03.0´n 19°02.0´w; map 4). large island in the danske øer group in the south part of jokelbugten. so named by the 1938–39 mørkefjord expedition because of its size. (storöen.) straight river 70ø (70°28.7´n 23°09.3´w). minor, straight river in south jameson land, so named by hermann aldinger during the 1931–34 treårsekspeditionen. stranda-huset 71ø (c. 71°52´n 22°45´w). norwegian hunting hut on the nw side of fleming inlet below sporfjeld, built by the møre expedition in august 1931. it is also known as flatstranda. the hut was swept away by a wave during a storm in 1953. stranddal 77ø-138 (77°04.0´n 23°12.9´w; map 4). valley in north dronning louise land containing strandelv, which drains bri tannia sø. named by the 1952–54 british north greenland expedition. strandelv 77ø-139 (77°00.0´n 23°02.4´w; map 4). river draining britannia sø in north dronning louise land. it follows the west margin of storstrømmen and at its north end flows over sandy beaches and terraces from which the name derives. the name was given by the 1952–54 british north greenland expedition, who had in 1951 named an ice-dammed lake on the same site adastra lake. strathclyde pynt 69ø-75 (69°43.6´n 23°36.0´w). cape where the west end of turner sund meets rømer fjord. named by malcolm slesser’s 1969 expedition for the university of strathclyde, to which the leader malcolm slesser was affiliated. strathclyde was the name given in the 9th and 10th centuries to a british kingdom which extended over the basin of the river clyde. it became a scottish province in the 11th century. (strathclyde point.) stratumbjerget 74ø-361 (74°26.9´n 20°09.2´w). mountain 679 m high in west wollaston forland, made up of alternating layers of sandstone and shale. so named during lauge koch’s 1938–38 expeditions by wolf maync and andreas vischer for the geological relationships. (stratumbjærget.) straumneset 72ø (72°43.6´n 22°44.2´w; fig. 14). elongate sand spit on south geographical society ø, ne of silja ø. so named on the nsiu maps of lacmann (1937) because it borders a sound with strong currents (= straum). straumpollen 72ø (72°43.5´n 22°40.5´w; fig. 14). west-facing bay on south geographical society ø, se of the scott keltie øer. used only on nsiu maps (lacmann 1937), the name was given for the strong currents (= straum). straumtangen – see strømtangen. strawberry peak 73ø (73°15.6´n 27°47.6´w). name used in a climbing report by odell (1943) for the 2268 m mountain north of lystergletscher, frænkel land. it was climbed by n.e. odell during the 1933 louise boyd expedition. the mountain, sometimes re ferred to as mount gore, has a summit composed of blood-red quartzite. stress-hytten 71ø (71°40´n 22°56´w). hut on the east side of fleming fjord, about 4 km from the head of the fjord. it was built by nordisk mineselskab in 1976 with material supplied by ‘stress tagelementer’ of fårevejle, sjælland. stribebjerg 73ø-704 (73°19.0´n 28°42.9´w). mountain 2565 m high in west frænkel land. named during lauge koch’s 1951 expedition by john haller and eduard wenk for its striped appearance, due to alternating granite and gneiss bands. stribedal 73ø-50a (73°58.6´n 21°21.2´w). minor valley on the north slope of stensiö plateau, nw hold with hope, draining west into blåelv. so named by eigil nielsen during the 1931–34 treårs eks pe di tionen because repetition of beds by faults produces a striped appearance. strindberg land 73ø-513 74ø-240a (73°50.0´n 25°00.0´w; maps 2, 4). land area bounded by geologfjord to the west and walters hausen gletscher and nordfjord to the east. named by a.g. nat horst in 1899 as strindbergs halfö after nils strindberg [1872– 1897], a swedish physicist who was one of the three lost members of andrée’s 1897 balloon expedition for whom nathorst’s expedition was searching. (strindbergs peninsula, strindbergs halvøya, strindberghalbinsel.) strindberg valley 73ø (73°45.8´n 24°48.8´w). name occasionally found used in norwegian reports for brogetdal in strindberg land, although the more common norwegian usage is giæverdalen or stordalen. (strindbergdalen.) strindberghuset 73ø (73°42.2´n 24°30.6´w). norwegian summer station built in june 1935 at the mouth of brogetdal, strindberg land, for salmon fishing. attempts at tinning salmon (arctic char) were made here in 1938. the station was renovated in 1954. it has also been known as laksehytta. (strindbergdalen, strindberg.) strindberghytta 73ø (73°42.2´n 24°30.6´w). norwegian hunting hut in strindberg land at the mouth of brogetdal, built by arktisk næringsdrift in 1930. in 1935 it was demolished, and the material used to build strindberghuset. (strindberghytten.) striped cliff 73ø (73°12.4´n 27°42.9´w). cliff at the bend of knæk dalen where the stream from lystergletscher meets knækelven. so named by louise boyd’s 1933 expedition because of the alternating dark and light layers of the banded gneisses. stripöya 73ø (73°58.2´n 21°08.5´w). island in the finsch øer group, so named on an nsiu map (1932a) for its long, narrow form. strittberg 71ø (71°56.5´n 23°35.3´w; map 5). peak about 1871 m high on the west side of spærregletscher. it was climbed, and so named, by the 1964 aac zürich expedition. strømbugt 70ø-414 (70°31.0´n 27°58.0´w). large bay on sw milne land, opposite the mouth of vestfjord. named during the 1967–72 ggu scoresby sund expeditions by svend funder for the marked tidal current along the coast of the bay. strømhytta – see strømnæshytten and villaen. strømhytta 73ø (c. 73°02´n 22°55´w). norwegian hunting hut on the north coast of geographical society ø, se of robertson ø. built by arktisk næringsdrift in august 1929, and named after ing wald strøm, one of the three hunters who built it. it was moved to sofia sund in august 1932 where it was known as stor-dalen. (ström hytta, strøm-hytten.) strømhytta 73ø (73°19.0´n 24°48.9´w). norwegian hunting hut on the north side of dusén fjord, sw of barrieren. built by arktisk næringsdrift in september 1930, 2 km east of the narrow part of the fjord which is subject to strong currents (= strøm). the hut has also been called dyrfaret (nsiu 1932c) and trangen. now a ruin. (strømmen, strømhytten.) strømmen 73ø-86 (73°54.6´n 21°54.2´w). narrow part of northern loch fyne marked by strong tidal currents, named by lauge koch’s 1929–30 expeditions. (strömmen, straumen.) strømmenhytten 73ø (73°53.2´n 21°52.5´w). danish hunting hut on the west coast of loch fyne, east hudson land, immediately south of strømmen. it was built by nanok in september 1950, and has also been known as danske villa. the norwegian hut on the east side of loch fyne has also gone under the similar name strømhytta, but is better known as villaen or norske villa. (strømshytta.) strømnæs 72ø-441 (72°42.5´n 26°47.0´w). peninsula half way 315 along röhss fjord, at the narrowest part where there is a strong tidal current. the name was used by eugène wegmann during the 1931–34 treårsekspeditionen. a ruined hut lies on the east side of the peninsula (see strømnæshytten). strømnæsdal 72ø-442 (72°41.1´n 26°50.5´w). valley in gletscher land draining into röhss fjord at strømnæs. so named by eugène wegmann during the 1931–34 treårsekspeditionen. strømnæshorn 72ø-442a (72°41.0´n 26°55.4´w). mountain in glet scher land between strømnæsdal and röhss fjord, so named during the 1931–34 treårsekspeditionen by eugène wegmann. strømnæshytten 72ø (72°42.4´n 26°47.7´w; fig. 82). norwegian hun ting hut on the south side of röhss fjord at strømnæs, built in july 1934 for arktisk næringsdrift. it was originally known as festningen. now a ruin. an old dog-sledge and a heavy wooden boat lie beside the hut, the latter abandoned here by anders busk in 1956 on the instructions of lauge koch. (strømhytten.) strømsbukta 76ø (76°14.6´n 20°01.5´w). norwegian hunting hut built in august 1933 by john giæver’s expedition 2–3 km west of kap peschel, ad. s. jensen land. strømsund 76ø-149 (76°41.1´n 21°26.1´w). narrow sound of the coast of se daniel bruun land. so named by j.p. koch’s 1912–13 expedition because they encountered a 4–5 knot strong current while negotiating the sound. (straumsund.) strømtangen 74ø-268 (74°01.2´n 22°01.4´w). low peninsula on the west side of the mouth of loch fyne. named on an nsiu map (1932a) as straumtangen, because it was built up by strong currents. stubba 73ø (73°22.8´n 22°09.1´w). river draining the southern giesecke bjerge, flowing in the present sindalen. so named on the 1932a nsiu map. the name in translation implies something short or stubby. stubbdalen 73ø (73°22.8´n 22°09.1´w). valley in the south giesecke bjerge, corresponding to sindalen, and carrying the river stubba. so named on the 1932a nsiu map. studer gletscher – see øvre studer gletscher, nedre studer glet scher. stuegulvet 71ø (71°51.3´n 25°05.6´w; map 5). low (1780 m) and easy summit on the north side of roslin gletscher. ascended on ski by the 1996 norwegian stauning alper expedition, it was situated north of one of their depots. ‘stuegulvet’ is a term used for a surface ‘as smooth as a dance floor’ stufenberg – see terrassebjerg. stugunosa 73ø (73°30.9´n 21°35.9´w). hill 252 m high north of mygg bukta. so named on an nsiu map (1932a; fig. 13), for its proximity to myggbukta station (stugu = house). stuttgarter spids 71ø (71°50.5´n 25°20.6´w; map 5). mountain on the south side of the head of roslin gletscher. climbed by karl herligkoffer’s expedition on 21 august 1966, and named after the south german city of stuttgart, capital of baden-würtemberg. styggbreen 74ø (74°13.3´n 22°32.4´w). lobe of wordie gletscher between scotstounhill and manley bjerg. used on the nsiu maps of lacmann (1937), the name was given for its grim and dangerous appearance (stygg = nasty). støvdal 77ø-132 (77°07.0´n 24°00.0´w; fig. 21). valley between the snouts of admiralty gletscher and britannia gletscher, filled by moraine, fluvial and aeolian deposits. so named by the 1952–54 british north greenland expedition because of the frequent dust spirals seen here during the summer, which gave rise to the name støvdal or dust bowl used in expedition accounts. due to the subsequent advance of britannia gletscher the site of the valley is now an ice-dammed lake. støvfanget 70ø-448 (70°21.0´n 29°44.7´w). sheltered area on the sw side of paul stern land between the glacier and the cliff, where large quantities of mica-dust collect (støv = dust). named by adrian phillips during the 1967–72 ggu scoresby sund expeditions. südprofil 74ø (74°43.4´n 20°02.6´w). geological reference locality on se kuhn ø, used by maync (1947) in his description of work during lauge koch’s 1938–38 expeditions. suess land 72ø-44 (72°59.0´n 26°20.0´w; maps 3, 4). land area bounded by kejser franz joseph fjord to the north and kempe fjord and dickson fjord to the south. named by a.g. nathorst in 1899 after eduard suess [1831–1914], an influential austrian geologist who was professor of geology at vienna from 1861. nathorst had translated a book by eduard suess into swedish. (suessland.) sugar basin 71ø (71°52.0n 25°31.1´w; map 5). name given in re ports of james clarkson’s 1961 expedition to the upper broad basin of spærregletscher. it is an area without crevasses which was named for the snow conditions. suhm bjerg 73ø-337 (73°26.4´n 22°11.7´w). mountain in the central giesecke bjerge. the name was proposed by the place name committee in 1939 to replace a suggestion by wolf maync and andreas vischer. it commemorates peter frederik suhm [1728–98], a danish historian who produced a 14 volume ‘historie af danmark’. håkampen has also been used. (suhms bjerg.) sukces gletscher 72ø-310 (72°00.4´n 23°58.3´w; map 5). glacier in the north werner bjerge. named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk, probably for the find here of a boulder in the moraine containing molybdenum. sukkertoppen 73ø-576 (73°53.4´n 29°24.3´w). nunatak west of j.l. mowinckel land, so named by arne høygaard and martin fig. 82. the ruin of the norwegian hunting hut, strømnæshytten, on the south side of röhss fjord at strømnæs. it was built in july 1934. 316 mehren in 1931 for its pyramid-like shape (sukkertoppen = the sugar loaf ). sulebak 72ø (c. 72°24´n 25°49´w). norwegian hunting hut on the south side of forsblad fjord, 2 km south of caledonia ø, said to have been built in 1931 for the møre expedition by o. åmbak and peder sulebak. it was also recorded under the name caledonia hytten, although in fact the projected hut was never built (p.s. mik kelsen 1994, 2008). sulugpik – see suluppik. sulugssût, sulugssûtikajik – see sulussuut, sulussuutikajik. sulugssûtikajîp kiámut kangertiva – see sulussuutikajiip kiammut kan g er tiva. sulugssûtikajîp orqungmut kangertiva – see sulussuutikajiip oqqum mut kangertiva. suluppik 70ø-180 (70°35.5´n 22°26.0´w). point on the east coast of hurry inlet between dumbrava and castor elv. recorded by the 1955 geodætisk institut name registration, the name translates roughly as ‘the place where one had diarrhoea’. (sulugpik.) sulussuut 71ø-220 (71°04.3´n 25°26.8´w). elongate island in the bjørne øer group with a prominent knife-edge ridge at the nw end. the greenlandic name, recorded by the 1955 geodætisk in sti tut name registration, was given for its appearance, translating as ‘the dorsal fin’. (sulugssût.) sulussuutikajiip kiammut kangertiva 69ø-53 (69°56.0´n 22°52.0´w). fjord on the north blosseville kyst. the name was recorded by the 1955 geodætisk institut name registration, and translates as ‘the fjord to the north of sulussuutikajik’. (sulugssûti kajîp kiámut kangertiva.) sulussuutikajiip oqqummut kangertiva 69ø-55 (69°55.0´n 23°00.0´w). fjord on the north blosseville kyst. the name was re corded by the 1955 geodætisk institut name registration, and derives from it being a ‘sheltered fjord behind sulussuutikajik’. (su lugs sûtikajîp orqungmut kangertiva.) sulussuutikajik [steward ø] 69ø-4 (69°54.3´n 22°52.0´w). island sw of kap brewster, north blosseville kyst. the name was recorded by the 1955 geodætisk institut name registration, and derives presumably from its shape as it translates as ‘the little dorsal-fin’. four substantial hunting houses have been built in a bay on the nw side of the island, and hunting families from scoresbysund have periodically over-wintered here since 1971 – see tsulitsuuligai. tuborg & sandell (1999) use the variation sulussugutikajik for the island. (sulugssûtikajik, sulusjuligæi, sulugssugetetajik, sulussuguti gajik.) sumpdalen 74ø-343 (74°33.8´n 19°35.1´w). wide depression in wollaston forland between albrecht bugt and falske bugt, so named during the 1936–38 two-year expedition by wolf maync and andreas vischer, because of its boggy nature (maync 1947). sun valley camp 72ø (c. 72°08´n 24°40´w). camp site on bersærker bræ in the north stauning alper, just below its junction with dunottar gletscher. the site was first used by malcolm slesser’s 1958 expedition. according to bennet (1972) it has become one of the most popular of climbers’ camp sites in the stauning alper. (sunshine corner.) sûnínguâ, sûnínguai, sûnínguakajik – see suuninnguaa, suuninng uai, suuninnguakajik. sunderland gletscher 77ø-127 (77°06.0´n 24°48.6´w; map 4). glacier in nw dronning louise land. the name was given by the 1952–54 british north greenland expedition for the sunderland flying boats of the raf which flew the expedition and equipment to britannia sø from zackenberg bugt. sunnmøresheimen 72ø (72°25.0´n 24°33.8´w). norwegian hunting station 2 km se of kap peterséns built by the møre expedition in 1930, and named after the sunmøre area in norway from which the expedition came. the station was also called vardevakt, but is more commonly known for its location as kapp petersens, or kap peter séns. sunnmøre had long traditions in arctic fishing, whaling and hunting, and was better equiped and more active than other areas of norway. (sunnmørs-heimen, sunnmoers-heimen.) sunnmøresheimen 72ø (72°53.8´n 25°43.9´w). original name used for the norwegian hunting hut at lumskebugten, south suess land, built in september 1934 by arktisk næringsdrift. the name was changed to mineralbukta because the kapp peterséns hunting station was at that time known by the same name (p.s. mikkelsen 1994). (sunnmørsheimen.) sunnmøresterrenget 74ø (c. 74°15´n 19°333´w). name used in some accounts of norwegian hunting activities for that part of wolla ston forland between kap borlase warren and herschellhus (south of herschell bjerg), where the first norwegian expeditions from the sunnmøre region over-wintered in 1908–09. sunshine corner – see sun valley camp. suomi bjerg 72ø-427 (72°44.9´n 26°50.0´w; map 4). mountain in ne gletscherland. so named by eugène wegmann and heinrich bütler, who climbed the mountain on 14 august 1933, for the fin nish members of the 1931–34 treårsekspeditionen. surprise elv 73ø-718, 74ø-202 (74°00.9´n 22°17.3´w). river in ne hudson land, named during the 1931–34 treårsekspeditionen by gunnar säve-söderbergh as surprise river. surøje-hytten 71ø (71°52.2´n 22°45.6´w). norwegian hunting hut on the nw side of fleming fjord, built by otto lapstun in sep tember 1954 for hermann andresen’s expeditions. it is now usually known as lapstun hytten. the original name was given because the stove smoked badly giving rise to smarting eyes (= surøje). it has also been known as søndre biot, fladestrand and fleming fjord hytten. susan’s peak 72ø (72°06.2´n 24°54.7´w, map 5). peak 2238 m high on the ridge south of major passet, stauning alper. climbed by the 1996 scottish mountaineering club expedition. susannetop 73ø-684 (73°29.8´n 27°04.7´w; map 4). mountain in sw andrée land, on the ne side of isfjord. named during lauge koch’s 1949–51 expeditions by john haller, after susanne hallerweisskopf (mrs john haller). the name was inspired by a light coloured s-shaped gneiss band on the west flank of the mountain. suselv 73ø-303 (73°52.4´n 22°01.4´w). river in east hudson land draining into loch fyne. the name was proposed by the place name committee to replace an unsuitable suggestion by helge g. backlund. sussex fjeld 71ø-360 (71°58.0´n 25°08.5´w; map 5). peak 2300 m high sw of sidney fjeld, stauning alper. the two peaks were named by the 1963 cambridge university expedition after sidney sussex college, cambridge, established on the site of a franciscan convent under the will of frances sidney, countess of sussex. both peaks were climbed on 3 august 1963. (sussex.) suuninnguaa 71ø-231 (71°06.3´n 22°35.1´w). hill on the floor of the upper part of klitdal, between jameson land and liverpool land. recorded by the 1955 geodætisk institut name registration, the name translates as ‘its little head’. (sûnínguâ.) suuninnguaa [sydkap] 71ø-34 (71°17.3´n 25°04.5´w). southfacing cape on the north side of the mouth of nordvestfjord. the name was recorded by the 1955 geodætisk institut name registration, and means ‘its little head’. (sûnínguâ.) suuninnguai 70ø-330 (70°25.0´n 21°50.6´w). peninsula between kap tobin and kap swainson, south liverpool land. recorded by the 1955 geodætisk institut name registration, the name translates as ‘the little head’. (sûnínguai.) suuninnguakajik 70ø-336 (70°25.9´n 21°43.3´w). cape a little ne of kap swainson, south liverpool land. one of the names recorded by the 1955 geodætisk institut name registration, it translates as ‘the bad head’. (sûnínguakajik.) suzanne bræ 77ø-43 (77°19.5´n 24°22.5´w; map 4). glacier be tween ymer nunatak and north dronning louise land, so named by the 1909–12 alabama expedition. girl’s name. (suzanne gla cier.) 317 suzanne nunatak 77ø-120 (77°20.9´n 24°03.6´w; map 4; fig. 21). small nunatak in suzanne bræ, north of dronning louise land, connected by a moraine to the east end of ymer nunatak. named by the 1952–54 british north greenland expedition. svalbardr 70ø (c. 70°17´n 23°00´w). the icelandic annaler report the discovery in 1194 of svalbardr or svalbarda í hafsbotn, the country of the cold coasts. there is disagreement as to the interpretation of the sailing instructions found in landnámabók, olaf tryggvasons saga, hauksbók and ivar baardsson’s account. danish authorities have generally argued that svalbardr is identical with the scoresby sund region (ryder 1892; rafn 1845; holm 1926), wheras norwegian opinion prefers an identification with spits bergen (e.g. tornøe 1944). svalbard is today the official group name for the five islands of spitsbergen proper, and four other islands of which bjørnøya is the southernmost. they were placed under the sovereignty of norway by the treaty of paris in 1920. (svalbarde, svalbardi.) svampebugt 74ø-186 (74°09.0´n 21°31.3´w). open bay west of granatelv on sw clavering ø. derived from the reference locality soppbukta used in nsiu botanical reports, named for the fungi or ‘sopp’ (psalliota) (svampe = sopp = fungus). (sopbugt.) svampebugthytten 74ø (74°09.3´n 21°31.4´w). norwegian hut at the mouth of granatdal, east of svampebugt, built in august 1926 by the 1926–28 foldvik expedition. it has also been known as granat hytta and sandviken. originally located 5 km farther east, it was moved to the present site in july 1927. svanning bjerg 73ø-339 (73°25.5´n 22°18.1´w). mountain in the cen tral giesecke bjerge. the name was proposed by the place name committee in 1939 to replace suggestions by wolf maync and andreas vischer, and commemorates the danish historian, hans swaning [1500–84]. it corresponds to skrubbtind and skrubbfjellet of norwegian maps. (svannings bjerg.) svartetua 72ø (72°50.8´n 21°59.4´w). peninsula or cape on east geographical society ø on the north side of cambridge bugt. so named on the nsiu maps of lacmann (1937) for the colour (svart = black). svarthammerhytta 75ø (75°57.9´n 20°48.2´w). norwegian hunting hut on the south side of bessel fjord, built in september 1932 by john giæver’s expedition. it has also been known as fredhaug. svarthamrane 72ø (72°53.5´n 23°04.6´w). mountain ridge on cen tral geographical society ø corresponding to the present jule kagen. used only on nsiu maps (lacmann 1937), the name was given for the black cliffs. svarva 74ø (74°09.4´n 21°20.1´w). stream on south clavering ø adjacent to sveiva. so named on nsiu maps of lacmann (1937) for the turbulence (= svarva) of the stream. svedenborg 72ø (72°59.0´n 24°33.4´w). norwegian hunting hut at the cape nw of svedenborg bjerg on west geographical society ø. built by arktisk næringsdrift in september 1930, it has also been known as joplassen, valborghytten and røvballehytten. svedenborg bjerg 72ø-48 (72°56.7´n 24°27.2´w; map 4). moun tain range on west geographical society ø. named by a.g. nat horst in 1899 as svedenborgs berg, possibly after gustaf wilhelm emanuel svedenborg [b. 1869], a swedish officer who was a reserve member of andrée’s 1897 balloon expedition. nsiu maps restrict usage to the sw peak only (lacmann 1937). (svedenborgfjellet, sveden borg bjærg, mt. swedenborg.) sveiva 74ø (74°09.3´n 21°19.3´w). stream on south clavering ø ad jacent to svarva. used only on nsiu maps (lacmann 1937), the name derives from the norwegian dialect word for a stream that curves. it appears on danish maps to be a branch of svarva. svejstrup bjerg 74ø-377 (74°40.2´n 21°16.3´w). mountain on the south side of the mouth of svejstrup dal. named by the 1948 leeds university expedition led by w.r.b. battle. (svejstrups bjerg, svej strups mountain.) svejstrup dal 74ø-179 (74°45.0´n 21°24.0´w; map 4). valley at the head of lindemann fjord between th. thomsen land and a.p. olsen land. this was one of the names which first appeared on the 1932 edition of the geodætisk institut 1:1 million scale map, which derives from lauge koch’s aerial observations in 1932 during the 1931–34 treårsekspeditionen. it was said to have been named after a head of department in grønlands styrelse. material for a norwegian hunting hut was landed at the mouth of the valley in august 1938, but it was never built. (svejstrups dal.) svejstrupdalshytten 74ø (74°39.9´n 21°03.1´w). danish hunting hut about 5 km up svejstrup dal, a.p. olsen land, built by nanok in may 1947. (svejstrup dal hytten.) svendsby 74ø (74°38.5´n 20°49.5´w). norwegian hunting hut on the south side of lindeman fjord, immediately west of lindeman fjord hytten. it was built in august 1932 by sigurd tøllofsen’s ex pe di tion, and has also been known as stordalshytten. svenskenæs 73ø-716 (73°14.2´n 26°21.4´w; map 4). eastern cape of frænkel land, where isfjord meets kejser franz joseph fjord. named by john haller and eduard wenk following explorations during lauge koch’s 1951 expedition. svensnæs 74ø (74°09.5´n 20°18.5´w). cape on east clavering ø, north of dahl skær. the name appears only on a sketch map in gustav thostrup’s 1921 logbook (møller 1939). sverdrup hytte – see solveigs hytta. sverdrupsnes 74ø (c. 74°21´n 19°11´w). name apparently used by the 1908–09 floren expedition for a cape north of kap borlase war ren, possibly the east flank of clark bjerg. it was presumably named after otto neumann sverdrup [1854–1930], norwegian commander of the fram during the drift across the arctic ocean from 1893 to 1895 led by fridthof nansen, and leader of his own expedition to the canadian arctic islands, also in the fram, from 1898 to 1902. p.s. mikkelsen (1994) indicates that sverdrupsnes has also been used for the norwegian hunting station borganes at kap borlase warren. sverresborg 72ø (72°50.9´n 22°56.8´w). norwegian hunting station on the south side of geographical society ø, east of tværdal, built by arktisk næringsdrift in 1929. named after sverre sørensen, who with søren richter and thor halle constructed the station. possibly also named after two castles of the same name in trond heim and berg, norway, built by the norwegian king, sverre. this outer coastal region is subject to heavy winter snow, gave poor hunting, and the station was abandoned in 1932. now a ruin. (gåse hytten.) svingnæs 77ø-118 (77°01.6´n 20°23.9´w). cape on the west side of southern sælsøen where the lake makes a pronounced swing to the west. named by the 1938–39 mørkefjord expedition, probably as a result of the exploration journey made by paul gelting and carlos ziebell in june 1939. svinhufvud bjerge 72ø-138 (72°26.3´n 23°35.2´w; maps 4, 5; fig. 78). mountain range on the sw side of traill ø with summits be tween 1000 m and 1380 m high. the name came into use during lauge koch’s geological expeditions in the 1930s, and is attributed to finnish geologists. it commemorates pehr evind svin hufvud [1861–1944], president of finland from 1931 to 1937. svinta 74ø (74°08.4´n 20°37.5´w). small stream on se clavering ø, the present fossilelv. used on the nsiu maps of lacmann (1937), the name derives from the norwegian dialect word for ‘fresh’. swiss peak 72ø (72°02.3 n 24°25.4 w; map 5). mountain 1769 m high nw of skelpas, north werner bjerge. bennet (1972) reports it was climbed by gerold styger in 1950, with the second ascent in 1958 by k. bryan and donald bennet. sydbjergene 70ø (70°06.0´n 23°17.9´w). descriptive name applied by some members of the 1924–25 scoresbysund colonisation expedition to the mountain ranges on the south side of scoresby sund, corresponding approximately to the present klinten (bengtsson 1927). sydbræ 70ø-70 (70°06.4´n 26°20.9´w; maps 3, 4). large glacier on 318 the south side of gåsefjord, which flows from south to north. named by carl ryder’s 1891–92 expedition as syd bræ because it lay due south of their winter harbour on danmark ø. the ams maps use sydgletscher. sydelv 72ø-104 (72°28.3´n 25°22.8´w). river in the south half of polhem dal draining south into forsblad fjord. named during the 1931–34 treårsekspeditionen by ove simonsen. sydelv 77ø-117 (77°05.6´n 20°40.2´w; map 4). river on the north side of sælsøen, notable for its very deep gorge. named during the 1938–39 mørkefjord expedition, probably by paul gelting and alwin pedersen. sydfjorden 70ø (70°10.0´n 27°15.0´w). name used for the present gåsefjord in ragnvald knudsen’s diaries of carl ryder’s 1891–92 expedition to the scoresby sund region. sydgavlen 78ø-22 (77°57.5´n 19°26.5´w). island east of hagen ø, named by the 1938–39 mørkefjord expedition which deposited depots here. it is the last large island in the row of islands south of hammeren, and has a triangular south face resembling the gable of a temple. ggu’s new topographic maps place the island entirely south of latitude 78°n. sydgletscher 71ø-296 (71°58.5´n 26°24.0´w; map 4). glacier on the south side of the west end of furesø, nathorst land. named during the 1954–55 lauge koch expeditions by hans zweifel for its n–s trend. sydhytten 73ø (73°27.5´n 20°53.7´w). danish hunting hut on the south side of hold with hope, west of kap broer ruys. built by nanok in august 1945, the hut is also known as kap broer ruys syd. sydhøjen 80ø-120 (80°09.5´n 22°24.5´w; map 4; fig. 24). penin sula on the north side of centrumsø with inuit ruins. named during operation groundhog 1960. sydkap [suuninnguaa] 71ø-34 (71°17.3´n 25°04.5´w; maps 3, 4). prominent south-facing peninsula between the mouth of nord vest fjord and nordøstbugt. named by carl ryder’s 1891–92 expedition as syd cap. hunters from scoresbysund spent long periods here from about 1934, with great success, and the ruins of their houses are found west of the cape. a more substantial house and store-house were built at the cape in 1946 by a danish telegraphist and his greenlandic wife with a view to fishing for salmon and shrimps, a venture abandoned after a year. some reports say his wife found it too lonely. hunters still occasionally spent periods at sydkap. see also kangertertivarmiit [sydkap]. (syd kap.) sydkap 78ø-39 (78°40.3´n 19°24.0´w; maps 1, 4). south cape of schnauder ø, jøkelbugten, named by the 1938–39 mørkefjord ex pedition. sydkronen 71ø-407 (71°48.6´n 23°36.0´w). mountain 1140 m high in the south part of the bjergkronerne massif, north of ørsted dal. named by katharina perch-nielsen during the 1967–72 ggu scoresby sund expeditions. sydkærene 74ø (74°27.9´n 20°34.1´w). locality south of zacken berg forskningsstation. the name is used as a reference locality in reports by visiting scientists. sydlige fligelyhytten 74ø (74°45.2´n 20°37.0´w). danish hunting hut north of the mouth of lindeman fjord, about 10 km south of blåbærdal, built by nanok in august 1931. it is also known as linde mannhytten. sydlige gneisnæs 76ø-159 (76°12.5´n 18°33.2´w; map 4). south ern of two gneiss ridges bounding areas of sediments on the east side of store koldewey. named by the 1906–08 danmark-ekspedi tionen as south gneiss naze. (südliche gneisnaes.) sydlige jægersundhytte 76ø (76°15.0´n 20°24.5´w). danish hunting hut built by nanok in september 1938 on the south point of nanok ø in the southern part of jægersund. it is officially known as has se ris hytten. (jægersundhytten.) sydney tinde 71ø (71°55.3´n 25°43.2´w; map 5). mountain on the east side of prinsessegletscher at the head of castor gletscher. named and first climbed by the 1967 berchtesdgaden expedition. (sidney tinde.) sydvejen 80ø-85 (79°58.4´n 20°35.0´w; map 4). branch valley on the south side of eastern rivieradal. named during lauge koch’s 1952–53 expeditions by erdhart fränkl. on new topographic maps the valley is entirely south of latitude 80°n. sydvestgletscher 72ø-317 (72°11.2´n 25°42.3´w). glacier on the sw side of schaffhauserdalen. named by john haller following explorations during lauge koch’s 1954 expedition. sydvestpynten 73ø-18 (73°35.2´n 23°58.8´w). cape on west gauss halvø, facing sw. the name was suggested by the place name committee in 1935, and has been variously placed on the rounded coastline. it is said to correspond to the original position of koldewey’s 1869–70 cap gauss (see also gauss halvø). a hunting hut about 5 km north of the point, sometimes known as sydvest pynten, is usually known as huttetu. sylbugten 76ø-280 (76°25.0´n 20°49.8´w). bay on the east side of godfred hansen ø, north of the mountain sylen. named by the 1938–39 mørkefjord expedition. sylen 76ø-48 (76°23.8´n 20°48.6´w). mountain on se godfred hansen ø. so named by the 1906–08 danmark-ekspeditionen, because its prominent pointed summit resembled an awl (= sylen). staff at danmarkshavn weather station in the 1950s referred to the mountain as jennovs næse (= jennov’s nose). sylfjeldene 70ø-236 (70°46.8´n 21°46.0´w). mountain ridge be tween horsens fjord and vejle fjord on the east coast of south liverpool land. so named by laurits bruhn during the 1931–34 treårsekspeditionen for its pointed summits (syl = awl). syltekrukken 76ø (76°51.0´n 18°47.4´w). landing strip and hut north of danmarkshavn weather station. knud ‘sylte’ nielsen and his younger brother bendt ‘lille sylte’ nielsen prepared a 300 m long airstrip and built a small hut in 1961, both being improved in 1964. the present hut was built in 1966, and in 1979 the old hut was moved to a new site where it is known as germania land hytten. the nicknames of the nielsen brothers derive from an incident with one of their dogs known as ‘syltetøj’ (= marmelade, jam) (steinert 1973). the landing strip was superceded in 1992 by a new strip built beside the weather station at danmarkshavn. syltoppene 72ø-24a (c. 72°20´n 24°33´w; maps 4, 5; fig. 78). range of spiked mountains in the ne stauning alper, named by a.g. nathorst in 1899 as syltopparne because of their needle-like summits. five of the summits including ochsenberg were climbed by the 1964 aac zürich expedition. (needle points, syltoppane.) sylva maria tinde 72ø-423 (72°56.7´n 26°42.9´w). mountain in south suess land, named during the 1931–34 treårsekspeditionen by eugène wegmann in the form sylva-maria massif. sylva-maria is a common girl’s name in switzerland. synna 73ø (73°41.9´n 22°04.8´w). river in east hudson land, a minor tributary to storelv, which flows south from nordhoek bjerg and synshovd. so named on an nsiu map (1932a), possibly for a river of the same name in the oppland region of norway. synshovd 73ø (73°44.9´n 22°01.8´w). mountain 1317 m high south of nordhoek bjerg, on the west side of loch fyne. so named on the nsiu (1932a) map, probably for its situation at the head of the river synna. syttendemajfjorden 76ø-45 (76°15.0´n 21°01.3´w). narrow fjord in northern ad. s. jensen land. the name was first used by the 1906–08 danmark-ekspeditionen in the form syttendemaj fjord. 17 may is the norwegian national day, and two members of the expedition were norwegian. skyggefjord has been used for the same feature. (17. maj fjorden.) syttenkilometernæshytten 76ø-194 (76°49.3´n 18°17.2´w). danish hunting hut north of the peninsula syttenkilometernæsset on the east coast of germania land. built by nanok in 1935, it is now a ruin. it was replaced by a hut built by danmarkshavn weather station personnel in march 1979, known as syttenkilometer næsset. 319 syttenkilometernæsset 76ø-40 (76°49.2´n 18°17.8´w; map 4). peninsula ne of danmarkshavn. so named by the 1906–08 dan mark-ekspeditionen in this form because of its approximate sledging distance from their base at danmark havn. there is an inuit settlement with 16 houses here. (17-kilometernæsset, seventeenkilometer naze.) syveren 71ø (c. 71°46´n 22°57´w). norwegian hunting hut built in 1954 by otto lapstun on the north side of fleming fjord, for helge ingstad’s expedition. all the wall elements of the hut were marked with the number seven. it has also been known as mellem-huset and funkis. syvstjernen 76ø-138 (c. 76°33´n 26°32´w; fig. 21). group of seven small nunataks in sw dronning louise land, so named by j.p. koch’s 1912–13 expedition (syvstjernen = the seven stars). syvsøstre bræ 71ø-443 (71°17.8´n 27°37.5´w; map 4; fig. 83). glacier on the west side of renland draining into edvard bay dal. so named by geoffrey halliday during the 1971 northern univer sities expedition, because the glacier is formed by the confluence of seven glaciers. sæfaxi dal 80ø (80°09.7´n 20°40.9´w). name occasionally used by fränkl (1954) for the valley west of marmorvigen in which sæfaxi elv runs. sæfaxi elv 80ø-75 (80°09.7´n 20°40.9´w; map 4; fig. 24). river draining from centrumsø to hekla sund. so called after the ice landic catalina ‘sæfaxi’, which made the first landing on cen trumsø on 31 july 1952. the name first appeared in the report by fränkl (1954) (sæfaxi = sea horse). sælhunden 79ø (79°23.6´n 19°32.9´w). small island off the ne coast of lambert land. the name was used by the 1996 myliuserichsen’s mindeekspedition, and was probably given for its shape. sælsøgletscher 77ø-98 (77°05.7´n 22°00.0´w; map 4). name originally used in some 1906–08 danmark-ekspeditionen reports for the glacier at the head of sælsøen. it was later proposed as a formal name by eigil knuth. (sælsö-gletscher.) sælsøen 77ø-22 (77°04.5´n 20°50.4´w; maps 2, 4). lake 30 km long forming the north boundary of daniel bruun land. the lake surface is about 4 m above sea level. so named by the 1906–08 danmark-ekspeditionen because on one occasion they saw what appeared to be a seal swimming near the outlet stream (trolle 1909). other expedition reports note the same origin for the name, but also say that the sighting was a mistake (thostrup 1911; koch 1916). trolle reports that the original name for the lake was store sø. lakse sø has also been used. (sælsø, seal lake, sælsöen, sæl lake, selvatn.) sælsøhytten 77ø-81 (77°02.5´n 20°16.4´w). danish hunting hut on the ne side of sælsøen, built by nanok in october 1933 at tvilling næs. now a ruin. it has also been known as tvillingnæshytten. sæmingfjellet 73ø (73°10.1´n 23°56.8´w). mountain 1625 m high on se ymer ø. so named on the 1932a nsiu map. 4. sænkning 76ø-265 (76°17.9´n 18°37.4´w; map 4). locality on the east side of store koldewey, used by the 1906–08 danmarkekspeditionen as a geological reference locality. sætherheia 72ø (72°46.3´n 22°10.0´w). part of the east flank of freycinet bjerg in se geographical society ø. so named on the nsiu maps of lacmann (1937) after carl sigmund sæther [1880– 1947], a norwegian who was british consul in tromsø from 1923, and was agent for british expeditions operating in the arctic. sætherhytten 76ø (c. 76°04´n 20°03´w). norwegian hunting hut built in september 1932 by john giæver’s expedition 2–4 km east of the mouth of trumsdal, now a ruin. it was named after carl s. sæther (see sætherheia). søbjergene 70ø-163 (70°47.0´n 22°16.4´w). mountain ridge be tween kalkdal and sødal, south liverpool land. so named by laurits bruhn during the 1931–34 treårsekspeditionen because of the numerous lakes. (søbjærgene.) sødal 70ø-167 (70°44.1´n 22°18.5´w; map 4). valley in liverpool land on the east side of hurry inlet. so named by laurits bruhn during the 1931–34 treårsekspeditionen because of the presence of several large lakes. (södal.) sødal 74ø-332 (74°07.5´n 23°59.7´w; map 4). valley in ole rømer land north of krumme langsø. named during lauge koch’s 1938–38 expeditions by heinrich bütler for the lakes in the valley. sødalen 80ø-45 (80°33.8´n 21°08.0´w; map 4). valley running west from the interior of ingolf fjord. so named by eigil nielsen during the 1938–39 mørkefjord expedition, because he thought the valley drained a large lake he named troldsøen. there are no lakes in the valley, and eigil nielsen appears to have misidentified an area of flat-lying snow as a lake. nielsen may also have been misled by lauge koch’s maps of the region drawn from the air in 1933, which show two large lakes (romer sø and centrumsø) draining into the east end of the valley. in fact centrumsø drains along a more southerly route into marmorvigen, while romer sø drains directly into ingolf fjord from the north. sødalshytten 74ø (74°31.5´n 20°59.7´w). danish hunting hut on the north side of the largest lake in sødal, behind zackenberg, built by nanok in june 1939. (søhytten, sødalhytten.) söderbergh plateau 74ø-231 (74°10.0´n 20°41.1´w). small plateau on se clavering ø west of moskusokseelv, named by lauge koch’s 1929–30 expeditions after gunnar säve-söderbergh [1910–48], who worked in this region in 1931–34 and 1936. he was a swedish palaeontologist noted for his work on devonian fossil fish, and the discovery of the stegocephalians. (søderbergh plateau, söderbergs plateau.) søelv [kaporniagaqarteq] 70ø-166 (70°43.9´n 22°24.2´w). river fig. 83. glacier draining west from the renland ice cap, named syvsøstre bræ for its seven tributaries. the john haller photograph collection, geus archive. syvsøstre bræ renland 320 draining the lakes in sødal, south liverpool land. named during the 1931–34 treårsekspeditionen by laurits bruhn. (söelv.) søgletscher 73ø-354 (73°54.6´n 25°02.8´w; map 4). glacier in central strindberg land, with an ice-dammed lake on its north side. named during lauge koch’s 1948–49 expeditions by hans r. katz. søgletscher 77ø-84 (77°12.2´n 20°43.8´w; map 4). glacier east of the south end of annekssøen. the name was suggested by the place name committee to replace a suggestion by the 1938–39 mørke fjord expedition. paul gelting and alwin pedersen had visited the area in may 1939. sølverbæk 81ø (81°05.9´n 13°18.5´w). river in ne kilen, kronprins christian land. the name is found on a coloured geological map of kilen printed in 1991 (pedersen 1991), and was named after ‘silver stream’, a locality in tolkien’s ‘lord of the rings’. sølvhorn 72ø-314 (72°17.8´n 24°52.1´w; map 5). mountain between linné gletscher and skjoldungebræ. climbed by peter braun and fritz schwarzenbach in july 1951, who named it silber horn (braun 1953). the name was approved in 1957 at the suggestion of john haller. it is a descriptive name for a horn-shaped sum mit with ice on all sides. (solvhorn.) sønderelv 70ø-54 (70°39.9´n 25°24.2´w; map 4). river sw of kap leslie, east milne land, draining south. named during the 1931– 34 treårsekspeditionen by hermann aldinger as südfluss. sønderelv 75ø-92 (75°50.2´n 19°45.9´w; map 4). river in east nørlund land, south of the mouth of bessel fjord. the name first appeared on a map in jennov (1939). a nanok hut on the north side of the river sometimes known as sønderelv-huset is officially known as hundehushytten. (søndre elv.) sønderelv 75ø (75°49.9´n 19°39.7´w). norwegian hut built in november 1933 on the south side of sønderelv for john giæver’s expedition. it was replaced in 1949 by astralhytten. sønderelv-huset – see hundehuset. sønderfjord 71ø-137 (71°03.6´n 21°53.4´w). fjord on the south side of storefjord, central liverpool land, so named during the 1931–34 treårsekspeditionen by laurits bruhn. søndergletscher 71ø-291 (71°55.1´n 23°48.8´w; maps 4, 5). gla cier in the south werner bjerge, flowing south. named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk. (söndergletscher.) sønderland 77ø-97 78-34a (77°45.0´n 21°53.0´w; map 4). south ern most part of hertugen af orléans land. named by the 1938–39 mørkefjord expedition. søndermarken 77ø-105 (77°22.1´n 21°03.8´w; maps 2, 4). trian gular area of land south of nordmarken, between annekssøen and valdemarsmuren. the name was apparently a modification of a suggestion by the 1938–39 mørkefjord expedition. sønderstrand 76ø-282 (76°22.9´n 20°55.3´w). flat coastal stretch of south godfred hansen ø. named during the 1938–39 mørke fjord expedition, perhaps by paul gelting who visited it in april 1939. søndersund 76ø-214 (76°22.5´n 20°57.0´w). sound south of god fred hansen ø, sw dove bugt. named by the 1938–39 mørkefjord expedition. sønderås 71ø-105 (71°37.4´n 22°17.6´w). ridge in south canning land. the name was first used by säve-söderbergh (1937) in the form southern ridge and derives from work during lauge koch’s 1936–38 expedition. søndre basisdal 71ø-103 (71°36.3´n 22°15.2´w). valley in se can ning land draining south to carlsberg fjord. the name appears to have first been used by säve-söderbergh (1937) in the form s. basis valley, and derives from surveying work during lauge koch’s 1936–38 expeditions. søndre biot 71ø (71°52.2´n 22°45.6´w). norwegian hunting hut on the nw side of fleming fjord about 10 km sw of kap biot. it was built in september 1954 for hermann andresen’s expedition, and has also been known as surøje, lapstun hytten, fleming fjord hytten and fladestrand. søndre gneissnæs – see sydlige gneissnæs. søndre mellemland 78ø-34 (78°08.0´n 21°36.0´w; maps 1, 2, 4). southernmost but one part of hertugen af orléans land, so named by the 1938–39 mørkefjord expedition. søndre muschelberg 75ø-52 (75°10.1´n 19°55.0´w). south-western of the two low mountains making up muschelberg, hoch stetter forland. so named during the 1931–34 treårs eks pe di tionen by hans frebold. (søndre muslingebjerg.) søndre muslingebjerg – see søndre muschelbjerg. søndre næs 76ø (76°45.3´n 18°39.3´w). name used occasionally in reports of the 1906–08 danmark-ekspeditionen for one of the peninsulas of danmark havn, probably østre havnenæs. søndre orienteringsø 76ø-254 (76°42.2´n 19°48.7´w). southern most island of the orienteringsøer in dove bugt. so named by the 1906–08 danmark-ekspeditionen. søplateauet 73ø-664 (73°43.9´n 25°24.5´w). plateau between morænedal and geologfjord, named by erdhart fränkl during lauge koch’s 1948–50 expeditions for the many small lakes. (sø plateau.) sørensenflya 72ø (72°49.9´n 22°49.5´w). hillside on geographical society ø, sloping down to vega sund ne of gåseøen (flya = plain). used only on nsiu maps (lacmann 1937), the name was given for sverre sørensen [b. 1899], a norwegian hunter who wintered in east greenland in 1929–31 and 1932–33. sørkjosen 72ø (72°41.7´n 22°08.4´w; fig. 14). bay on se geo graphical society ø, west of kap mcclintock. used only on nsiu maps (lacmann 1937), and named for its position relative to nordkjosen a bay to the nw.(sörkjosen.) søryggen 74ø-331 (74°06.3´n 23°47.9´w). ridge between two large lakes, vibeke sø and krumme langsø. named during lauge koch’s 1938–38 expeditions by heinrich bütler, originally in the form seerücken. søspidsen 74ø-46 (74°35.0´n 18°45.4´w). mountain 333 m high on east sabine ø. named by karl koldewey’s 1869–70 expedition as seespitze, apparently because the summit cliffs descend steeply to the sea below. søstersøer 77ø-54 (77°15.8´n 23°45.8´w; map 4). two very similar, adjacent lakes in north dronning louise land, named by the 1909–12 alabama expedition (søster = sister). (søstersøerne, söster söerne.) søstjernen 73ø-253 (73°01.2´n 22°18.5´w; map 4). island in the brochs øer group with a cross-like shape, originally named on the nsiu (1932a) map as korstrollet. both danish and norwegian names translate as ‘starfish’. søstrene 70ø-350 (70°06.0´n 22°21.2´w). three similar mountain peaks about 1070 m high on a ridge on northern savoia halvø. so named by laurits bruhn during the 1931–34 treårsekspeditionen (søstrene = sisters). søvngængerbjerg 70ø-422 (70°40.3´n 29°04.4´w). mountain 1790 m high north of rolige bræ. so named by laurent jemelin during the 1967–72 ggu scoresby sund expeditions, apparently because the party reached the summit at the end of a long and exhausting day, almost like sleep-walkers (= søvngænger). t t-sø 71ø-324 (71°41.8´n 27°04.9´w; map 4). large lake shaped like the letter ‘t’ on the north side of inner nordvestfjord. the name came into use during the 1950s when the lake was used as a landing site for catalina aeroplanes attached to lauge koch’s expeditions. tafelbjerg 74ø-44 (74°38.7´n 18°47.3´w). mountain 428 m high on ne sabine ø. named by karl koldewey’s 1869–70 expedition as tafelberg, for its flat top, and possibly also for a mountain of similar name in austria. (mt tafelberg.) 321 tagbjergene 74ø-325 (74°00.6´n 23°18.1´w; map 4). range of mountains on the south side of promenadedal, north hudson land. named during lauge koch’s 1936–38 expeditions by hein rich bütler (tag = roof ). taget 71ø-270 (71°57.0´n 24°01.8´w; map 5). mountain in the cen tral werner bjerge, on the north side of sirius gletscher. named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk, and climbed by bearth in 1953 (taget = the roof ). taggletscher 74ø-198 (74°11.0´n 21°10.2´w; map 4). ice cap on south clavering ø, due north of eskimonæs station. the name was first used during the 1931–34 treårsekspeditionen by gelting (1934), and was given for the roof-like appearance. tagstenstop 70ø-445 (70°12.2´n 29°28.9´w). mountain 1360 m high on a nunatak on the se side of vestfjord gletscher. so named by w.e. adrian phillips during the 1967–72 ggu scoresby sund expeditions because the rocks weather into cleaved slabs suitable for roofing slates (= tagsten). tait bjerg 71ø-26 (71°29.2´n 22°36.9´w; map 4). mountain 710 m high on the west side of carlsberg fjord. named as cape tait by william scoresby jr. in 1822, probably after william tait [1793– 1864], bookseller and publisher, and a well known figure in the social life of edinburgh. scoresby’s cape was found subsequently to be a mountain and the name changed accordingly. takkerne 72ø-168 (72°28.7´n 21°59.4´w). mountain or cape on east traill ø, north of kap parry. so named during lauge koch’s 1936–38 expeditions by hans p. schaub because of its serrated summits. takkerne 73ø-328 (73°55.7´n 22°36.3´w; map 4). mountain range in ne hudson land, on the north side of stordal. adapted from a suggestion by heinrich bütler arising from his work during the 1936–38 two-year expedition. tancredia river 70ø (70°30.7´n 22°37.2´w). name used by harris (1931) for the river flowing in tancrediakløft. tancrediakløft 70ø-284a (70°30.7´n 22°37.2´w). conspicuous ravine in neill klinter on the west side of hurry inlet. named by alfred rosenkrantz during lauge koch’s 1926–27 expeditions as tancredia kløft for the occurrence of fossil lamellibranchs. (tan cre dia klöft.) tanden 76ø-225 (76°55.5´n 21°20.4´w). mountain on the south side of inner mørkefjord. so named by the 1938–39 mørkefjord expedition led by ebbe munck and eigil knuth, for the shape (tand = tooth). tandlaegetinde 71ø (71°57.1´n 25°04.6´w). summit 2350 m high in the upper reaches of sefström gletscher, stauning alper. climbed by the 2001 scottish mountaineering club expedition. the leader of the expedition, colwyn jones, is a dental surgeon (= tandlæge). tangebugt 70ø-249 (70°56.3´n 21°41.3´w). bay on the east coast of liverpool land, on the north side of kap greg (tang = seaweed). tangen 74ø-121 (74°21.3´n 21°50.8´w). large delta on the west side of clavering ø almost blocking the fjord apart from the channel revet. named by lauge koch’s 1929–30 expeditions. muskos öyra is used on lacmann’s (1937) maps. tangodden 76ø-292 (76°55.6´n 20°20.8´w). headland west of mørke fjord station, west of gamma havn. named by the 1938–39 mørkefjord expedition led by ebbe munck and eigil knuth, probably for the occurrence of seaweed. tantallon gletscher 72ø-358 (72°01.5´n 25°11.7´w; map 5). narrow glacier with tantallon spids at its head, on the ne side of sefström gletscher, north stauning alper. named tantallon gla cier by malcolm slesser’s 1958 expedition. tantallon spids 72ø-360 (72°02.2´n 25°07.2´w; map 5). rock peak 2480 m high with many spires on the ne side of sefström gletscher, north stauning alper. first climbed by malcolm slesser’s 1958 expedition, and named after tantallon castle, east lothian, a douglas stronghold dating from c. 1375. (tantallon.) tantalus 71ø (71°46.3´n 25°18.7´w; map 5). mountain 2477 m high on the ne side of orion gletscher, south stauning alper. first climbed by james clarkson’s 1961 expedition, and named after the tantalus of greek mythology. tappedal 71ø-423 (71°06.7´n 27°42.1´w; map 4). valley draining tappesø in sw renland. so named by j.d. friderichsen during the 1967–72 ggu scoresby sund expeditions because the icedammed lake tappesø drains (= tappe) through the valley. tappesø 71ø-422 (71°10.0´n 27°46.7´w; map 4). lake in sw renland, ne of rypefjord, at the margin of eielson gletscher. so named by j.d. friderichsen during the 1967–72 ggu scoresby sund expeditions because the lake drains (= tappe) through tap pe dal. taraxacumfjeld 71ø-392 (71°02.0´n 23°00.0´w). summit 1261 m high sw of pothorst bjerge, north jameson land. the name was proposed by russel marris following his journeys in the region in 1968, and given for one of the 25 greenland species of dandelion. tartaajik [glasgow ø] 70ø-235 (70°48.7´n 21°39.1´w). island off the east coast of liverpool land. recorded by the 1955 geodætisk institut name registration, the name means ‘it looks like a seal’s kidney’. (tartâjik.) tartâjik – see tartaajik. taseq 71ø (71°26.5´n 25°14.5´w). name used occasionally for one of the lakes of holger danske briller north of sydkap (taseq = the lake). taseq qúteq – see taseq qutteq. taseq qutteq 70ø-200 (70°30.3´n 21°54.7´w). lake ne of scores by sund town. one of the names recorded by the 1955 geodætisk institut name registration, it translates as ‘the uppermost lake’. (taseq qúteq.) taskedalen 74ø (c. 74°16´n 19°23´w). name used by the 1908–09 floren expedition for one of the valleys west of kap borlase warren. position uncertain. tassiusark 76ø (76°45.9´n 18°39.4´w). name used during the 1906– 08 danmark-ekspeditionen for the eastern bay of danmark havn (poulsen 1991). tattaalakajia 70ø-342 (70°03.6´n 22°45.1´w). moraine ridge on roma gletscher, volquaart boon kyst. recorded by the 1955 geo dætisk institut name registration, the name translates as ‘the little stone ridge’. (tavtâlakajia.) taurobjerg 71ø (71°37.6´n 24°59.1´w; map 5). mountain 1860 m high at the head of leo gletscher, south stauning alper. so named by the 1970 university of dundee expedition, because two of the climbers who made the first ascent were born under the zodiac sign taurus, and the name fitted with others in the vicinity. taurus glacier 71ø (71°43.6´n 25°24.8´w). tributary to orion glet scher, south stauning alper, so named by james clarkson’s 1961 expedition for the constellation of the zodiac. tavlen 74ø-386 (74°04.9´n 29°01.0´w; map 4). nunatak in northern hobbs land, originally named tafelbjerg for the flat-lying basalts by hans r. katz during lauge koch’s 1951 expedition; the name was changed to tavlen (= the board) by the place name com mittee. the highest point at 2400 m was climbed by katz. tavtâlakajia – see tattaalakajia. taymors fjell 74ø (74°33.9´n 19°18.2´w). name occasionally used by norwegian hunters for falkebjerg, wollaston forland, and also for the hunting hut at its foot known as falkberget or falske bugt hytten. tear drop lake 70ø (70°51.1´n 23°35.3´w). small oval lake where the 1989 greenland milne land expedition camped during their climbing expedition. tectonic valley 76ø (76°24.4´n 19°00.0´w). name used informally by bronner (1948) in his geology report of louise boyd’s 1938 expedition for a t-shaped transverse valley on the west side of store koldewey. a major fault zone was found here. teddys udkig 74ø-2a (74°32.8´n 18°48.9´w). name given by eske bruun in 1971 to the south slope of germania bjerg, sabine ø, to 322 commemorate louis rostock-jensen [1899–1966], affectionately known to his colleagues as ‘teddy’. in 1923 he was second mate on the ship teddy, and had climbed the slope twice a day to examine ice-conditions prior to leaving the east greenland coast on 9 august. the ship was lost in the ice, and rostock-jensen took over leadership of the party and played a significant role in the rescue of the crew. he was subsequently promoted to commander, and was a director of baltica. teebjerg 70ø-110 (70°51.6´n 22°53.3´w). mountain in east jame son land west of the head of hurry inlet. named tee mt. by alfred rosenkrantz and tom harris during lauge koch’s 1926–27 expeditions because of a supposed resemblance to an over-sized golf tee. teglbjerg 73ø-109 (73°08.5´n 23°30.5´w). mountain on east ymer ø, named during the 1931–34 treårsekspeditionen by gunnar säve-söderbergh as mt. tegelberget, because the brick-red rocks were the colour of roof tiles. (brick mountain.) teichert bjerg 74ø (c. 74°36´n 23°04´w). mountain ne of marianne nunatakker, north of wordie gletscher. the name is only found on the sketch map by th. johansen published in koch (1940 fig. 34). the map was drawn during a sledge journey along the margin of the inland ice between wordie gletscher and bessel fjord by a party of four men during the 1931–34 treårsekspeditionen, a party which included the geologist curt teichert. teigandalen 72ø (72°58.8´n 22°59.1´w). valley on geographical society ø, draining ne, the present græsdal. the name is found in lacmann (1937), and was given for the clumps (= teigan) of grass. the name is also used on 1951 usaf aeronautical charts. tektonbjerget 74ø-362 (74°24.8´n 20°01.6´w). mountain in west wollaston forland, so named during lauge koch’s 1936–38 expeditions by wolf maync and andreas vischer because of the tectonic relationships. tellplatte 75ø-29 (75°01.2´n 18°25.5´w; map 4). low hill with pronounced plateau-like summit on south shannon. so named by karl koldewey’s 1869–70 expedition, probably for the tellsplatte by the vierwildstättersee in switzerland. wilhelm tell was a swiss legendary hero said to have died in 1354. tellplatte pynt 75ø (75°00.0´n 18°23.5´w peninsula 2 km ne of kap david gray, shannon, the se projection of the 300 m high tell platte. the name is used in den grønlandske lods (1968). teltdammen 74ø (74°28.0´n 20°34.7´w). reference locality south of zackenberg forskningsstation. the name has been used by visiting scientists. teltskær 76ø (76°41.7´n 18°32.4´w). name used by c.s. poulsen (lund bye 1984) during the 1906–08 danmark-ekspeditionen for a skerry off kap bismarck, se of danmark havn, probably the pre sent måtten. tennes 74ø (74°19.2´n 21°52.9´w). small peninsular south of revet on the west coast of clavering ø. so named on the nsiu maps of lacmann (1937) for the place of the same name in balsford in the troms district of norway. it was the home of meyer olsens and hans olsens, two hunters of the 1926–28 foldvik expedition. tennholm 73ø (73°28.0´n 21°30.9´w). name used on an nsiu map (1932a) for a small island in mackenzie bugt, the site of a tern colony. the same island had been called ternøya in 1900 by gustav kolthoff for the same reasons. the grønlandske lods (1968) uses ternholmen. (tärnholmen, tern island.) tennskj. 73ø (73°03.2´n 22°37.5´w). small island in the broch øer group, so named on an nsiu map (1932a) for the terns. (tennsky.) tent peak 71ø (71°38.7´n 25°17.2´w; map 5). snow-capped peak at the head of jupiter gletscher, east of wedge peak, south stauning alper. first climbed by james clarkson’s 1961 expedition, and so named because it was capped by a neat gable of snow resembling a tent. termier gletscher 71ø-253 (71°57.6´n 23°46.5´w; map 5). glacier in the east werner bjerge, draining ne into the head of kolledal. named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk, perhaps after henri-françois-émile termier [1897–1989], a noted french geologist known mainly for his work in morocco. termografengen 76ø-293 (76°55.8´n 20°20.2´w). area west of mørke fjord station where paul gelting undertook experiments with soil thermometers during the 1938–39 mørkefjord expedition led by ebbe munck and eigil knuth. termometerfjeldet 76ø-186 (76°46.6´n 18°38.5´w). hill 138 m high north of danmark havn. so named by the 1906–08 danmarkekspeditionen because meteorological instruments including ther mometers were placed here by alfred wegener immediately after arrival. (thermometerfjeld, termometer fjeld, thermometer hill, ther mo meter mountain.) termometersøen 76ø (76°47.0´n 18°39.5´w). lake north of termo meterfjeldet. the name was used in the published diaries of the 1906–08 danmark-ekspeditionen (poulsen 1991; thostrup 2007; j. løve, personal communication 2009). tern island 72ø (72°14.4´n 23°47.8´w). name used by university of dundee expeditions between 1968 and 1974 for a small island in the mouth of noret. terneskær 73ø-46 (73°56.3´n 20°55.5´w). small islands off the north coast of hold with hope near the mouth of rødelv, named by lauge koch’s 1929–30 expeditions in the form terne skerries for the colonies of arctic terns. (tennungane.) terneskæret 76ø-76 (76°48.0´n 19°05.1´w). small island in the north part of stormbugt. so named by the 1906–08 danmarkekspeditionen for the arctic terns which nested abundantly here, and also on many other small islands and skerries. (sea-swallow skerry, tern reef, terne skerries.) ternevigen 70ø-417 (70°57.2´n 28°06.8´w). bay on the north side of harefjord, where there are many arctic terns. the name was given during the 1967–72 ggu scoresby sund expeditions. terningen 72ø (72°40.7´n 21°56.1´w). small island off the coast of se geographical society ø at kap mcclintock. named for its quadratic shape (terning = dice), and for the norwegian sealer terningen of tromsø which brought a norwegian hunting expedition to east greenland in 1928. terrassebjerg 74ø-42 (74°38.3´n 18°28.1´w). mountain 426 m high on lille pendulum. named by karl koldewey’s 1869–70 expedition as stufenberg, probably for the step-like profile of successive basalt lava flows. the sw flank of the mountain projecting into the sea has been called kap stufenberg (e.g. by den grønlandske lods 1968). the hut at the foot of the mountain was built by the 1928 hird expedition. terrassefjeld 77ø-115 (77°06.6´n 21°08.8´w). mountain on the north side of the inner part of sælsøen. the name was adapted by the place name committee from a suggestion by the 1938–39 mørkefjord expedition led by ebbe munck and eigil knuth, and derives from nearby terrassekløft. terrassehytten 75ø (75°50.2´n 19°40.2´w). danish hunting hut built by nanok in may 1931 on the north side of sønderelv, wollaston forland. it is also known as hundehuset and sønderelv-huset. (ter rasso hytten.) terrassekløft 77ø-89 (77°06.5´n 20°53.4´w; map 4). ravine on the north side of sælsøen, just west of midternæs. there are terraces in its lower part, on one of which a danish hunting hut was built. named by the 1938–39 mørkefjord expedition led by ebbe munck and eigil knuth. terrasseodde 70ø-82 (70°19.5´n 24°49.6´w; map 4). peninsula on the east side of terrassevig, volquaart boon kyst, so named during the 1931–34 treårsekspeditionen by laurits bruhn for the ter races, formed by nearly horizontal basalt lava flows. terrassepynt 71ø-39 (71°05.3´n 27°44.6´w). peninsula in the inner part of rypefjord, so named by carl ryder’s 1891–92 expedition. the expedition camped on a terrace here during their first sledge journey in april 1892. 323 terrassesøerne 76ø-299 (76°56.8´n 20°15.4´w). five small lakes on the terraces between mørkefjord station and the south end of sælsøen. named by the 1938–39 mørkefjord expedition led by ebbe munck and eigil knuth. terrassevig 70ø-81 (70°18.9´n 24°51.1´w; map 4). bay on vol quaart boon kyst adjacent to terrasseodde. named during the 1931–34 treårsekspeditionen by laurits bruhn. terrassohytten 75ø (75°50.2´n 19°40.2´w). danish hunting hut on the south side of the mouth of bessel fjord, also known as hunde huset. it was built by nanok in 1931. terre de france 77ø, 78ø (78°00´n 21°50´w). original name pro posed by the duke of orléans in 1905 for a newly discovered land area, which he wished named after his homeland france. he reluctantly agreed to the request of the danish administration to change it to terre du duc d'orléans, the present hertugen af orléans land. teufelcape island 76ø (76°23.3´n 20°24.5´w). name used by am drup (1913) for djævleøen in dove bugt, the island of which teu fel kap is the east cape. teufelkap 76ø-5 (76°23.0´n 20°09.8´w; maps 2, 4). eastern cape of djævleøen in the sw part of dove bugt. named as teufelscap by karl koldewey’s 1869–70 expedition for its sinister appearance, seen first as an imposing reddish wall through the fog in april 1870. several subsequent travellers have commented on the emi nent suitability of the name. (teufels cap, teufel cape, devil’s cape.) teufelsschloss 73ø-504 (73°22.2´n 25°29.3´w; map 4; figs 84, 85). isolated mountain 1340 m high on the coast of se andrée land. so named by karl koldewey’s 1869–70 expedition, because it resembled a colossal ruined castle. curt teichert, who mapped the region in 1931, recorded there was nothing ‘devilish’ about the mountain (unpublished report, geus archive). it was first climbed by noel e. odell and walter a. wood during the 1933 louise boyd expedition. the next recorded ascent was by erdhardt fränkl and fritz schwarzenbach in 1950. (djäfvulsslottet, devil’s castle.) tevla 73ø (73°31.8´n 20°33.2´w). minor tributary of the river glom men in se hold with hope. so named on an nsiu map (1932a). th. sørensen land 71ø-442 (71°20.5´n 28°18.0´w; map 4; fig. 41). land area between flyverfjord and edvard bay dal. named by geoffrey halliday following botanical work during the 1971 northern universities expedition, for thorvald julius sørensen [1902–73], a danish botanist who had published the botanical work of the 1931–34 treårsekspeditionen together with gunnar seidenfaden. sørensen was professor of botany at the university of copenhagen from 1956 to 1972, and director of the botanical gardens and museum. th. thomsen land 74ø-180 75ø-83a (74°52.0´n 21°26.0´w; maps 2, 4; fig. 15). land area bordered by grandjean fjord, svejstrup dal, tvegegletscher and fligely fjord. mapped by lauge koch during flights in 1932 on the 1931–34 treårsekspeditionen expedition, it was named after thomas thomsen [1870–1941], danish ethnographer and curator at the national museum. he was a mem ber of the expedition committee. (th. thomsens land.) thala vig 70ø (70°25.5´n 21°55.2´w). the name is used in den grønlandske lods (1968) for the bay adjacent to kap tobin known as uunarterajiip kangerterajiva. the ice-strengthened cargo and expedition ship thala dan regularly visited scoresbysund and kap tobin on summer supply voyages. built in 1957 by the j. lauritzen shipping company as a polar expedition ship, the thala dan made many voyages to the arctic and antarctic. it was sold to the brazilian navy in 1982, renamed barao de teffé, and scrap ped in rio de janeiro in 2007. the great claw 72ø (72°07.4´n 25°22.8´w). hanging glacier on the on the east side of alpefjord, north of gully gletscher. the name was used informally by boyd (1935). the great snow crest – see great snow crest. the highway 70ø (70°28.8´n 23°23.2´w). name used by herman aldinger (1935) during the 1931–34 treårsekspeditionen for the long, low terrace running parallel to the coast of south jameson land and now known as flakkerhuk. the island 76ø (76°46.1´n 18°40.2´w). name appearing in some reports of the 1906–08 danmark-ekspeditionen for an area just west of the mouth of østerelven, danmark havn, which has the appearance of an island during the melt. the name was considered unsuitable and not approved. the rock finger 72ø (72°10.6´n 24°40.5´w). prominent minor peak on the south side of harlech gletscher, north of poplar, north stauning alper. first climbed by the 1963 imperial college expedition, it was named for its appearance. (the finger.) theodolite hills 70ø-a3 (70°27.6´n 23°12.6´w). minor range of hills in south jameson land west of the mouth of raukelv. so named by herman aldingerer (1935) during the 1931–34 treårsekspedi tionen, probably because the hills were used during surveying. theodolitplateau 74ø-120 (74°20.2´n 21°30.1´w; map 4). nw plateau area of clavering ø rising to about 700 m. named by lauge koch’s 1929–30 expeditions in the form theodolit plateau because oskar kulling began a series of theodolite measurements here. (theodolithögda.) theodolitskær 77ø-64 (77°25.8´n 19°46.0´w). small, rocky island south of joinville ø in skærfjorden. so named by david malmquist during the 1931–34 treårsekspeditionen, presumably because theo dolite measurements were made here. (theodolitskærene.) theresabjerg 72ø-381 (72°01.8´n 23°25.6´w). mountain in north scoresby land, on the west side of majdal. named by hans kapp during the 1957–58 lauge koch expeditions. girl’s name. third river 72ø (72°33.4´n 24°05.7´w). name used by university of dundee expeditions between 1968 and 1974 for a minor stream draining into holm bugt, sw traill ø. thomas bjerg 71ø-289 (71°51.9´n 24°05.5´w; map 5). summit on the ridge between aldebaren gletscher and breithorn gletscher, south werner bjerge. named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk, probably for a friend, although officially it was said to be for the ‘thomas process’ in mineralisation. thomas thomsen næs 77ø-18 (77°13.7´n 18°14.3´w; map 4). peninsula on the ne coast of germania land with large inuit ruins. so named by the 1906–08 danmark-ekspeditionen after thomas marius thomsen [1870–1941], a danish archaeologist who had as sisted in the preparation of a report on the expedition’s finds (thostrup 1911). thomsen was inspector at the national museum, copenhagen, from 1919, and was noted for several important excavations in denmark. a hut was built here in 1938 by willy knutsen for the 1938–39 norsk–franske polarekspedisjon. on some maps (e.g. usaf charts) the name is placed against the more conspicuous cape 7 km further north. (thomas thomsen’s nose.) thomson klippe 76ø-308 (76°59.6´n 25°06.3´w; map 4). cliff on the north side of admiralty gletscher, nw dronning louise land. one of the names given by the 1952–54 british north greenland expedition for notable scientists, it commemorates the british phy sicist sir joseph john thomson [1856–1940], chiefly known for his discovery of the electron. thora ø 72ø-332 (72°42.1´n 22°50.9´w). small island in vega sund. name proposed by søkortarkivet in 1956–57 following sur veying of the channel through vega sund as an alternative ap proach for ships en route to nyhavn. it was given for the thora dan, a 5050-ton polar ship built in 1956 for the j. lauritzen shipping company, which sailed mainly in greenland and finnish waters. thorkild vogts hytta, thorolf vogts hytta – see vogt-hytta. thornøestua – see tornøstua. thorolf vogts bugt 72ø (72°42.0´n 22°16.5´w). name occasionally used by norwegian hunters for a small bay on geographical society ø where the richter-hytta was built in september 1929. 324 thorshanesø 70ø-432 (70°29.8´n 27°47.7´w). small lake on sw milne land. named during the 1967–72 ggu scoresby sund expeditions by max fumasoli for the numerous grey phalarope (= thorshane) observed here. thorshanesø 76ø-243 (76°48.9´n 19°08.0´w). small lake on winge kyst in south germania land. so named by the 1906–08 danmarkekspeditionen for the grey phalarope, observed to be breeding in this and other small lakes in 1907 and 1908. thorsten ø 79ø (79°18.8´n 19°08.7´w). small island off ne lambert land, the present panoramaø. the name was used by the 1996 mylius-erichsen’s mindeekspedition, who misplaced their pano rama ø northwards to the present gamle jim øer. thorstensenvika 72ø (72°48.6´n 22°10.8´w). innermost part of cambridge bugt in east geographical society ø. so named on the nsiu maps of lacmann (1937) for john thorstensen [b. 1907], a norwegian telegraphist who manned the myggbukta radio station in 1932–33. thors cafe 77ø (77°32.1´n 19°08.0´w). hut built in the spring of 1966 for slædepatruljen sirius about 3 km north of kap amélie, stormlandet. it stands side-by-side with kap amélie hytte. thors hammer sø 74ø (74°30.2′n 20°37.8′w). small lake in the area known as morænebakkerne, north of zackenberg forsknings sta tion. the name is used as a reference locality by scientists studying lake ecosystems. three sisters 73ø (73°40.7´n 25°56.3´w). series of three closely spaced summit towers on the north side of grejsdalen, andrée land. climbed by the 2007 army boreal zenith expedition. threms pynt 70ø (c. 70°31´n 26°48´w). point on the north side of fig. 84. the 1340 m high mountain teufelsschloss on the coast of south-east andrée land that karl koldewey’s 1869–70 expedition likened to a colossal ruined castle. from: verein für die deutsche nordpolarfahrt in bremen (1873–74). fig. 85. the distinctive 1340 m high mountain teufelsschloss on the north side of central kejser franz joseph fjord. john haller photograph collection, geus archive. 325 føhnfjord, where carl ryder’s 1891–92 expedition left christoffer threms to look after their boat on 12 august 1891. the name is used in helge vedel’s diary of the expedition (gulløv 1991). thun søerne 72ø-456 (72°58.7´n 26°37.5´w; map 4). group of lakes in west-central suess land, so named during the 1931–34 treårsekspeditionen by eugéne wegmann for the swiss lakes of the same name. wegmann explored the region in august 1933. thurweiser kopf 71ø (71°53.7´n 25°39.6´w; map 5). mountain on the west side of spærregletscher between hecate glacier and pollux glacier. named and first climbed by the 1967 berchtesgaden expedition. thyre spids 72ø-198 (72°12´n 23°58´w). minor prominence on the east side of rungsted elv, north scoresby land. named by prospecting teams associated with lauge koch’s 1948–49 expeditions for thyra or tyre danebod [d. 935], wife of the danish king gorm den gamle. she is known only from inscriptions on two rune stones in jelling, and tradition associates her with the construction of the danevirke. (thyres spids.) tiber tinde 72ø (72°03.5´n 25°08.0´w). mountain 2460 m high in the stauning alper, sw of korsspids. climbed on 23 july by sandro pucci’s expedition, and named after the river tiber which flows through rome. tidevandsvigen 73ø (73°59.6´n 21°09.0´w). small bay on the se side of lille finsch. the name is used in den grønlandske lods (1968). tidselbjerg 74ø-408 (74°01.6´n 22°35.7´w). mountain south of wor die bugt, north hudson land. the name is attributed to paul stern who worked with lauge koch’s expeditions from 1955 to 1958 (tidsel = thistle). tiedemannfjellet 74ø (74°26.3´n 21°12.5´w). mountain on north clavering ø. used only on nsiu maps (lacmann 1937), and named after j.l. tiedemanns tobaksfabric, oslo, which financed the 1932 nsiu aerial photography in east greenland. tillit nunatak 71ø-382 (71°54.2´n 29°44.0´w; map 4). nunatak in west charcot land, with small outcrops of the rock type tillite. named during the 1967–72 ggu scoresby sund expeditions. tillitekløft 73ø-563 (73°31.9´n 24°51.8´w). ravine in east andrée land, draining into geologfjord. named by christian poulsen during lauge koch’s 1929 expedition as tillite canyon, because of the occurrence of late precambrian glacial deposits (tillites). this locality is placed incorrectly on the official place name maps, and as a consequence published gi maps also give the wrong location (there are no tillites at the authorised location). hambrey & spen cer (1987) pointed out the error. (tillitkløft.) tillyrie 71ø (71°57.0´n 25°01.5´w; map 5). peak about 2415 m high in the upper reaches of sefström gletscher, stauning alper. climbed by the 1998 scottish mountaineering club expedition, the name means nipple. timeglasset 77ø-134 (77°06.0´n 23°27.0´w). hill with two sum mits in north dronning louise land. the name was given by the 1952–54 british north greenland expedition, possibly because the hill in plan has an hour-glass-like shape (timeglasset = the hourglass). tinderne 72ø-34 (72°27.1´n 25°51.7´w). range of mountain peaks on the north side of forsblad fjord. named tinnarne by a.g. nathorst’s 1899 expedition for the spiky summits. (the pinnacles, tindane.) tindernes dal 71ø (71°05.0´n 26°50.0´w). name used in a report by christian vibe in larsen (1960) and by andersen (1960), for a valley in renland, the present catalinadal. it was named for the high mountain peaks on both sides of the valley (tinderne = pinnacles). tintagel fjeld 72ø-493 (72°07.5´n 24°44.0´w; map 5). mountain about 1800 m high at the head of bersærkerbræ, north stauning alper. first climbed by john hunt’s 1960 expedition, and named tintagel for tintagel castle in cornwall. the castle, dating from 1150 and built on the site of a celtic monastry, was according to legend the birthplace of king arthur. tioram gletscher 72ø-357 (72°01.2´n 25°17.8´w; map 5). glacier in the north stauning alper, sw of sefström gletscher, named by john haller and malcolm slesser after nearby tioram spids. tioram spids 72ø-356 (72°01.3´n 25°20.2´w; map 5). mountain 1800 m high on the sw side of sefström gletscher, north stauning alper. first climbed by malcolm slesser’s 1958 expedition that named it for a nordic castle in west invernesshire. (tioram.) tirefour 72ø (72°02.0´n 25°07.7´w). rock tower 2140 m high on the north side of sefström gletscher, north stauning alper, climbed by graham tiso’s 1968 expedition. tita 73ø (73°31.6´n 20°35.6´w). minor tributary of the river glommen in se hold with hope. so named on an nsiu map (1932a), possibly for the river of the same name in the nordtrøndelag district of norway. titanitspids 72ø-309 (72°01.0´n 23°53.0´w). mountain in the werner bjerge, named by peter bearth and eduard wenk during lauge koch’s 1953–54 expeditions for the finds of titanium-bearing minerals. titlingen 72ø (72°40.7´n 22°42.2´w; fig. 14). small island in vega sund sw of kap hovgaard, the present anita ø. used only on nsiu maps (lacmann 1937), and named for its size (titling = tiddler, or small fish). tjældebjerget 69ø-183 (69°03.1´n 31°45.8´w). mountain in northern kong christian ix land, climbed and named by lawrence r. wager’s 1935–36 expedition as tilted mountain, because the cliffs of lavas appeared to be tilted. tobias dal 73ø-53 (73°45.4´n 21°00.0´w; map 4). major valley in hold with hope. in april 1927 lauge koch sent his greenlandic assistant, tobias gabrielsen, to investigate the valley to find a route from the outer coast to loch fyne. tobias otto mikael gabrielsen [1878–1945] was a west greenlander who participated in numerous expeditions, including the 1906–08 danmark-ekspeditionen, lauge koch’s 1926-27 expedition and alfred wegener’s last expedition on the inland ice in 1930. (tobias valley, tobiasdalen.) tobias gletscher 80ø-112 (80°46.0´n 17°29.5´w; map 4). glacier in the prinsesse elisabeth alper, draining se into ingolf fjord. named by john haller following explorations during lauge koch’s 1956–58 expeditions, probably after tobias gabrielsen, in tribute to his work on the 1906–08 danmark-ekspeditionen (see also tobias dal). tobias ø [tuppiap qeqertaa] 79ø (79°20.6´n 15°46.5´w; maps 1, 4). island about 2 km long and 1500 m across with an ice cap about 35 m high, situated in the greenland sea about 80 km from the coast of ne greenland. a number of associated small islets led to it originally receiving the name tobias øer. the new land was dis covered during a research cruise by the german ice-breaker polarstern in 1993, when the first landing was made by helicopter. on 28 april 2001 a landing was made by a ski-equipped twin otter and a single island was recorded (bennike et al. 2006, 2009). the island was named after the greenlander tobias gabrielsen; see also tobias dal. sightings of supposed land off the coast of ne greenland have periodically been made since 1907 by various early explorers: see fata morgana landet. (tobias øer.) tobiashytten 73ø (73°43.9´n 21°23.9´w). danish hunting hut in the upper part of tobias dal, hold with hope, built by nanok in au gust 1938. tommelen 70ø-440 (70°29.3´n 29°08.4´w). tongue of ice from rolige bræ extending southwards. named by laurent jemelin during the 1967–72 ggu scoresby sund expeditions for its association with djævlehånden and langemanden (tommelen = thumb). tommelen 76ø-43 (76°14.3´n 20°27.9´w; map 4). prominent south cape of tvillingerne, an island north of ad. s. jensen land. so named by the 1906–08 danmark-ekspeditionen because it is thumb-shaped. tomsborg 75ø (75°03.9´n 18°54.0´w). danish hunting hut on the 326 west coast of shannon, about 4 km north of kap tramnitz. built by nanok in september 1948, and named after hans thomsen who helped build it. it is also known as kap tramnitz hytten. toni kurz spids 71ø (71°56.4´n 25°40.1´w; map 5). mountain on the west side of spærregletscher, between castor glacier and pollux glacier. first climbed by the 1967 berchtesgaden expedition, who named it after the bavarian guide toni kurz [1913–1936] from berchtesgaden, who died on the north face of the eiger in tragic circumstances. torbern bergman bjerg 73ø-36 (73°45.6´n 23°48.9´w; map 4). mountain 1515 m high in moskusokselandet, north of the mouth of moskusoksefjord. named by a.g. nathorst’s 1899 expedition after torbern olof bergman [1735–1784], a swedish scientist noted for his contributions to qualitative and quantitative chemical analysis. (torbern bergmans berg, torbern bergman mountain, torben-bergmanberg, mt. torbern bergman.) toretind 74ø-276 (74°01.1´n 22°26.3´w; map 4). mountain in the east nørlund alper, ne hudson land. so named on the nsiu (1932a) map, and derived from an old norwegian personal name. torkjellfjellet 73ø (73°23.8´n 22°54.4´w). mountain on the south side of gauss halvø, corresponding to part of the hjelmbjergene. so named on the 1932a nsiu map, possibly for thorkel of herjolf ness, a norse farmer in greenland. tornøehøgda 72ø (72°46.7´n 22°34.4´w). mountain on south geo graphical society ø. so named on nsiu maps of lacmann (1937) after johannes kristoffer tornøe [b. 1892], a norwegian who took part in nsiu expeditions to svalbard and greenland. he was also secretary of nsiu. tornøestua 75ø (75°27.9´n 21°38.5´w). norwegian hunting hut built in august 1933 by john giæver’s expedition on the north side of smallefjord. named after j. k. tornøe, see above. it is also known as smallefjordhytten. (thornøe-hytten.) torteerniarfik 70ø-349 (70°06.8´n 22°21.1´w). hillside west of kap brewster on volquaart boon kyst. recorded by the 1955 geo dætisk institut name registration, the name translates as ‘where one collects rhododendrons’. (tortêrniarfik.) tortêrniarfik – see torteerniarfik. torv bræ 69ø (69°59.0´n 23°08.0´w). name used on 1951 usaf aeronautical charts for the major unnamed glacier that reaches the coast at steward ø. it lies south of the ice plateau officially known as torvgletscher. the variation torvbræ has appeared on several maps. torvet 70ø-268 (70°06.3´n 23°27.4´w). small glacier east of gavlen on volquaart boon kyst, surrounded on three sides by high mountain ridges. so named during the 1931–34 treårsekspeditionen by laurits bruhn for its appearance (torv = market place, square). torvgletscher 70ø-269 (70°04.6´n 23°16.4´w; maps 3, 4). large flat glacier or small ice cap on volquaart boon kyst forming the plateau above torvet. named during the 1931–34 treårsekspedi tionen by laurits bruhn. see also torv bræ. toscano gletscher 71ø-295 (71°57.3´n 26°38.9´w; map 4). glacier on the se side of frederiksdal near furesø, nathorst land. so named during lauge koch’s 1954–55 expeditions by hans zweifel, because eduard wenk smoked his last ‘toscano’ cigar of the sum mer here. tour carrée 72ø (72°13.5´n 25°03.3´w). peak 2250 m high in the north stauning alper at the head of frihedsgletscher. named and climbed by claude rey’s 1970 expedition. tour chartreuse 71ø (71°57.5´n 25°44.2´w; map 5). mountain 2372 m high at the head of glacier des tours, east of prinsesse gletscher. named and first climbed by claude rey’s 1968 expedition. tour vercors 71ø (71°57.8´n 25°43.4´w; map 5). mountain about 2520 m high at the head of glacier des tours, east of prinsesse gletscher. named and first climbed by claude rey’s 1968 expedition. tour de pavot 71ø (71°54.6´n 25°52.6´w). rock tower about 1750 m high on the west side of prinsessegletscher. named and first climbed by claude rey’s 1968 expedition. tours des camaïeux 71ø (71°50.5´n25°39.0´w; map 5). peak about 2500 m high at the head of prinsessegletscher, east of col de fure soe. named and first climbed by claude rey’s 1968 expedition. tove birkelund fjeld 81ø (81°15.8´n 13°54.1´w). hill in nw kilen, kronprins christian land. the name is found on a coloured geo logical map of kilen printed in 1991 (pedersen 1991), and was given for tove birkelund [1928–1986], professor of geology at the university of copenhagen from 1966. tovika 72ø (72°41.8´n 22°14.8´w; fig. 14). bay on south geo graphical society ø, divided into two parts by a flat sandy peninsula. used only on nsiu maps (lacmann 1937), and so named be cause it has two parts (to = two). trafsa 74ø (74°06.3´n 21°14.9´w). stream on south clavering ø, the present østerelv. only used on the nsiu maps of lacmann (1937). traill hytten 72ø (72°52.7´n 24°01.7´w). norwegian hunting hut built in august 1929 by arktisk næringsdrift in northern traill ø, on the south side of vega sund. it is also known as østhytten and snøheim. traill ø 72ø-2 (72°40´n 23°43´w; maps 3, 4; fig. 29). large island bounded by vega sund and kong oscar fjord. named traill island by william scoresby jr. in 1822 in compliment to a highly estee m ed friend, thomas stewart traill [1781–1862], who became pro fessor of medical jurisprudence at edinburgh university. scoresby had given the name to the southern of the two eastern peninsulas of the island, and it was a.g. nathorst who first used the name in its present sense. (trail ö, trail island, trailløya, traills ø, traill insel.) traill-iup immikkeertivi [haslum øer] 72ø-57 (72°27.9´n 24°05.5´w; maps 4, 5). island group off sw traill ø on the south side of holm bugt. the greenlandic name was recorded by the 1955 geodætisk institut name registration, and translates as ‘traill’s islands’. (traill-iup ingmikêrtive.) traill-iup ingmikêrtive – see traill-iup immikkeertivi. trammen 72ø (72°48.4´n 22°51.9´w). small island in central vega sund adjacent to gåseøen. used only on nsiu maps (lacmann 1937) and so named because the island is small with a step-like shape (trammen = small step). trangen 73ø (73°19.0´n 24°48.9´w). name occasionally used for the norwegian hut built in september 1930 for arktisk næringsdrift at the narrow part of the head of dusén fjord (trang = narrow, tight). it is more usually known as dyrfaret or strømhytten. trangfjorden 74ø (74°27.2´n 20°57.9´w). name used by norwegian hunters in the 1930s for the relatively narrow stretch of tyroler fjord on the north side of clavering ø, bounded to the north by the steep cliffs of zackenberg and to the south by those of the eiger. it was also used by danish hunters, and is found in some scientific reports. trangfjordhuset 74ø (c. 74°28´n 21°03´w). norwegian hunting hut on the north side of central tyrolerfjord, which they usually called trangfjorden. the hut was built in september 1927 by the foldvik expedition, and also goes under the names meyer-hus and zacken berghuset. trangsund 73ø (73°58.0´n 21°09.8´w). narrow sound between stille ø (kilöya) and stripöya in the finsch øer group. so named on the 1932a nsiu map. trangsund 76ø-173 (76°16.5´n 20°43.8´w; map 4). narrow sound between the island tvillingerne and the mainland to the south. so named by the 1938–39 mørkefjord expedition led by ebbe munck and eigil knuth (trang = narrow). zielers sund has also been used. trap sø 74ø (74°30.2′n 20°36.5′w). small lake in the area known as morænebakkerne, north of zackenberg forskningsstation. the name is used as a reference locality by scientists studying lake eco 327 systems. trappebjerg 73ø-711 (73°09.0´n 28°30.5´w). mountain 2500 m high on the west side of gregory gletscher, west frænkel land. so named by john haller and eduard wenk following explorations during lauge koch’s 1951 expedition, because the ascent over stratified rocks resembled climbing a staircase (= trappe). the name may have been intended for the lower 2350 m high summit to the ne. trappedal 70ø-423 (70°36.8´n 28°57.8´w). valley on the north side of rolige bræ with a profile resembling a staircase (= trappe). named by laurent jemelin during the 1967–72 ggu scoresby sund expeditions. trappesøer 71ø (71°43.1´n 22°31.8´w). name used for a locality on the east side of wegener halvø in some geological reports, where shallow drill cores were taken from upper permian rocks. traunsteiner spids 71ø (71°53.6´n 25°26.7´w). mountain 1950 m high on the ridge between duart gletscher and the upper basin of spærregletscher. it was climbed by karl herligkoffer’s 1966 expedition, and named after the small town of traunstein at the foot of the bavarian alps. tre søstre 70ø-380 (70°16.8´n 29°07.5´w). three closely similar adjacent peaks in western gåseland, on the north side of vind blæsedal. named during lauge koch’s 1958 expedition by eduard wenk. tredie hvide 74ø-171 (74°19.4´n 20°36.2´w). part of a mountain range on ne clavering ø, named by arne noe-nygaard and gunnar säve-söderbergh during the 1931–34 treårsekspedi tionen. originally three peaks were given the names erste weisse, zweite weisse and dritte weisse, for their colour. the names were subsequently used in danicised form as 1. hvide, 2. hvide and 3. hvide. see also første hvide and anden hvide. trefjord bjerg 71ø-350 (71°16.6´n 22°40.5´w; map 4). mountain in east jameson land west of inner carlsberg fjord. so named by john h. callomon during lauge koch’s 1958 expedition, because rivers draining its flanks flow into three different fjords, carlsberg fjord, hurry inlet and scoresby sund. trefoden 70ø-197 (70°36.7´n 21°54.5´w). mountain in south liverpool land, the culmination of three converging ridges (trefod = tripod). so named by laurits bruhn during the 1931–34 treårs eks pedi tionen. trefork gletscher 76ø-316 (76°52.5´n 24°37.5´w; map 4). glacier in dronning louise land draining into trefork sø. named by the 1952–54 british north greenland expedition. trefork sø 76ø-315 (76°55.8´n 24°17.5´w; map 4). large lake in central dronning louise land between admiralty gletscher and borg gletscher. the name was given by the 1952–54 british north greenland expedition for its shape, the three arms of the lake resembling an inverted ‘t’. tregletscherdal 72ø-395 (72°05.0´n 23°16.4´w). valley in north scoresby land, nw of antarctic havn, where three glaciers merge. named by hans kapp during lauge koch’s 1957–58 expedition. tregletscherfjeld 73ø (73°55.6´n 29°29.1´w). name used by eigil niel sen (1935) for an ice-capped mountain at the head of blåelv in northern hold with hope, drained by three glacier outlets. (mt. tre gletscherfjeld, trejøkelfjæld.) trekant 72ø (72°04.7´n 25°40.1´w). mountain 2250 m high at the head of trekantgletscher, west of alpefjord. climbed and so named by wolfgang weinzierl’s 1970 expedition. trekanten [immikkeertikajik kiattikajik] 71ø-119 (71°16.6´n 21°42.4´w; map 4). island on the east coast of liverpool land. so named by laurits bruhn during the 1931–34 treårsekspeditionen for its triangular (= trekant) shape. hildegard island has also been used. trekanten 74ø (74°02.0´n 21°59.1´w). mountain in the tag bjergene, north hudson land. used only on nsiu maps (lacmann 1937), and named for the pyramidal form (trekant = triangle). trekanten 76ø-323 (76°50.6´n 25°22.6´w; map 4; fig. 21). small nunatak in west dronning louise land, on the south side of borg gletscher. the name was given by the 1952–54 british north greenland expedition for its triangular shape viewed from the north. trekanten 76ø (76°51.5´n 19°37.1´w). north point of nørre orienteringsø with a triangular shape, a useful landmark during sledge journeys by staff at danmarkshavn weather station. the name was said to be in use by danish hunters, and is noted in jennov (1963) as the site of a barnacle goose colony. trekantgletscher 72ø-318 (72°07.8´n 25°35.8´w; maps 4, 5). gla cier on the west side of inner alpefjord. named during lauge koch’s 1954 expedition by john haller. trekroner 77ø-81a (77°00.1´n 20°12.1´w; map 4). steep and barren mountain 360 m high, east of the south end of sælsøen, noted for its colony of barnacle geese. the name was used as a reference locality in several of the 1906–08 danmark-ekspedi tionen reports, and is described as having three summits with valleys between. it was named after the fortress of the same name at the mouth of copenhagen harbour. (trekronerfjeldet.) trekronerhytten 77ø (c. 77°01´n 20°01´w). danish hunting hut east of trekroner, germania land. it was built by nanok in the spring of 1938. officially known as pashytten, it has also been known as schulzhytten, hvalsletten and slettehytten. it has now disappeared (p.s. mikkelsen 2008). treogtredivekilometernæsset – see fyrretyvekilometernæsset. trespids – see dreispitz. tresteinane 72ø (72°41.5´n 21°53.7´w). small skerries off se geo graphical society ø near kap mcclintock. so named on nsiu maps of lacmann (1937) because there are three rocks (= trestein ane). trestrømmen 71ø-246 (71°59.3´n 23°57.0´w). glacier in the wer ner bjerge on the east side of østre gletscher. named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk for the three branches of the glacier. treyarnon 72ø (72°08.0´n 24°55.2´w; map 5). pinnacle about 2700 m high on the ne ridge of hjørnespids, north stauning alper. named and climbed by the queen mary college expedition on 13 august 1968. trianglen 71ø-384 (71°48.5´n 27°01.8´w). triangle-shaped mountain 1330 m high in frederiksdal, south nathorst land. named during the 1967–72 ggu scoresby sund expeditions for the shape, and for the locality of the same name in copenhagen. trianglen 74ø (74°14.6´n 19°30´w). the name has been used by hun ters of østgrønlandske fangstkompagni for a triangularshaped delta sw of kap borlase warren. trias elv 71ø-186 (71°37.3´n 24°10.6´w). river draining gurre holm bjerge, and flowing into schuchert flod. named by hans stauber during lauge koch’s 1936–38 expeditions for the rocks of triassic age. kempter (1961) used triasdal for the valley in which the river flows. triasdal 71ø-135 (71°03.4´n 22°21.2´w; map 4). valley west of the head of storefjord, central liverpool land. so named by helge g. backlund during the 1931–34 treårsekspeditionen, for the occurrence of triassic rocks, which at this locality rest unconformably on crystalline rocks. triaskæden 72ø-227 (72°05.5´n 23°46.2´w; map 4). mountain ridge on the east side of the bay mesters vig. named by prospecting teams associated with lauge koch’s 1948–49 expeditions for the age of the rocks. trigoniaelv 70ø-281a (70°29.0´n 22°39.9´w). minor tributary to lakseelv nw of kap stewart in se jameson land. named by alfred rosenkrantz during lauge koch’s 1926–27 expeditions as trigonia elv, for the fossil lamellibranchs. trinity fjeld 71ø-363 (71°58.0´n 25°17.2´w; map 5). rock peak about 2800 m high on the divide between canta bræ and 328 krabbegletscher. climbed by the 1963 cambridge university ex pedition and named after trinity hall, cambridge, founded in 1350, and the only college to have retained the name ‘hall’. the name is slightly misplaced on published geodetic institute maps. (trinity.) trinity glacier 71ø (71°58.0´n 25°10.9´w). name used by the 1963 cambridge university expedition for a glacier east of trinity fjeld, stauning alper, the present canta bræ. trinity gletscher 74ø-376 (74°43.7´n 21°48.9´w). glacier in north a.p. olsen land flowing north to svejstrup dal. named by the 1948 leeds university expedition to acknowledge the help given to them by trinity college, cambridge. it was also the college of d.s. brock, one of the expedition members. trinity, founded in 1546, is the largest of the cambridge university colleges. (trinity glacier.) trinucleus 74ø-130 (74°14.0´n 20°48.8´w). mountain on se clav ering ø with three summits, all about 1480 m high. named as mt. trinucleus by lauge koch’s 1929–30 expeditions (seidenfaden 1931). see also monacleus and binucleus. trio grand 71ø (71°08.3´n 26°03.7´w). summit 2185 m high on the ridge nw of grundvigtskirken, renland. climbed and named by the 2007 west lancashire mountaineering group expedition. trip-trap-træsko 76ø-77 (76°44.6´n 19°03.5´w). line of three small islands or skerries north of store koldewey, so named by the 1906–08 danmark exspedition. the phrase is used in denmark for objects in descending or ascending order of height, e.g. wooden shoes. (trip-trap-træsko skerries, trip-trap-træsko islands.) triple lochs 72ø (72°13.4´n 23°58.3´w). name used by the univer sity of dundee expeditions between 1968 and 1974 for three small lakes on the lower slopes of domkirken, east of rungsted elv. triton glacier 71ø (71°38.0´n 25°25.5´w; map 5). small glacier in the south stauning alper, a tributary to løberen (neptune glacier). named by james clarkson’s 1961 expedition after the largest satellite of neptune, a name derived from the merman of greek mythology. tritonskaret 71ø (71°38.9´n 25°19.6´w; map 5). name used by the 1996 norwegian stauning alper expedition for the pass between triton glacier and canis minor glacier, where they camped at 1800 m altitude. tritontind 71ø (71°38.4´n 25°21.3´w; map 5). mountain about 2150 m high at the head of triton glacier. it was climbed and so named by the 1996 norwegian stauning alper expedition. troels-lund bjerg 73ø-336 (73°29.1´n 22°14.2´w). mountain in the central giesecke bjerge. the name was proposed by the place name committee in 1939 to replace suggestions by wolf maync and andreas vischer. it commemorates troels frederik troelslund [1840–1921], a danish historian noted for his monumental book ‘daglig liv i norden i det 16. århundrede’. the mountain corresponds to durin, dvalin and gunsteinsfjellet of norwegian maps. (troels-lunds bjerg.) trolddal 69ø-73 (69°48.0´n 23°31.0´w; map 4). valley west of turner ø on the blosseville kyst, used by malcolm slesser’s party in 1969 on their route to skottepasset and steno gletscher. slesser describes it as a beautiful valley with an eerie cirque at its head, a suitable abode for trolls (trold = ogre, troll). troldedal 75ø-66 (75°45.4´n 20°56.0´w). valley in nørlund land draining into the head of agnete sø. the name originated from the wintering party at kulhus during the 1931–34 treårsekspedi tionen, and was approved for many years in the form troldedalen. troldehaven 77ø-96 78ø-22a (77°56.0´n 18°44.0´w; maps 1, 2, 4). complex of islands south of storøen. named by the 1938–39 mørkefjord expedition led by ebbe munck and eigil knuth, and described on a spring sledge journey as resembling a fairy tale town of castles and houses of different styles and ages which would merit seven stars in a baedekar (pioneer travel guide produced by karl baedekar from 1827 onwards). troldmarkerne 74ø-365 (74°22.1´n 20°32.1´w). basalt plateau between dolomitdal and djævlekløften, ne clavering ø. so named during lauge koch’s 1936–38 expeditions by wolf maync and andreas vischer because of the peculiar weathering. troldsøen 73ø-157 (73°29.2´n 20°38.8´w). lake in se hold with hope, named on an nsiu map (1932a) in the form trollvatnet, probably for its mysterious or enchanted setting. there are numerous similar place names in norway. (lake troldsøen.) troldsøen 80ø-39 (c. 80°31´n 22°09´w). name used for a supposed lake in sødalen, west of ingolf fjord. named by eigil nielsen during the 1938–39 mørkefjord expedition led by ebbe munck and eigil knuth for the “large lake occupying sødalen or, if there are several lakes, the northeasternmost of them” (nielsen 1941 pp. 12–13). a large lake named troldsøen is shown on nielsen’s map, but he appears to have been misled by snow in the valley bottom as this lake does not exist, and there are no significant lakes in the valley. nielsen’s map suggests that he assumed the large lake known as centrumsø drained through this valley into the head of ingolf fjord, as shown on a lauge koch map (see e.g. drastrup 1945). however, centrumsø drains along a more southern route. trollebotne 69ø, 70ø (68°–70°n). this name appears in a latin inscription on a 1668–69 map by thord thorlacius (steenstrup 1886, 1889) against the coastal region from 68°–70°n, which indicates that this major embayment was so called by the old icelanders after the giants (trolls) which lived there. trompeteren bastion 79ø-40 (79°25.6´n 20°11.9´w; map 4). mountain in north lambert land. one of a group of five names given by the place name committee for dogs used on the 1906–08 danmark-ekspeditionen. they replaced names suggested by john haller. ‘trompeteren’ was a ragged, sorry looking, completely apa thetic dog, but would periodically get up, stick his nose in the air and howl. tromsdal 73ø (73°30.0´n 23°36.0´w). norwegian hunters name for the south end of paralleldal on gauss halvø, named by john giæver in 1930 for its resemblance to the area around tromsø. a hunting hut built at the mouth of the valley (73°30.4´n 23°40.2´w) was also known as tromsdalen, although more usually under the name dalheim. (tromsdalen, new tromsdal.) tromsdalstinden 73ø (73°31.3´n 23°22.6´w). norwegian hunters name for a mountain in paralleldal on gauss halvø, probably the west end of sederholm bjerg. named by john giæver in 1930 for its resemblance to the area around tromsø. tromsyra 73ø (73°59.2´n 21°59.3´w). peninsula about 5 km south of strømtangen. the name is used on lacmann’s (1937) maps. tromsøtind 72ø (72°06.4´n 24°58.5´w; map 5). peak about 2250 m high on the spiky ridge south of dansketinden, stauning alper. so named by the 1996 norwegian stauning alper expedition because the first ascent was made by two climbers from tromsø. trondfjellet 73ø (c. 73°22´n 22°31´w). mountain on southern gauss halvø, part of højsletten. so named on the 1932a nsiu map, possibly for the mountain of the same name in the troms district of norway. trugbjerg 73ø-432 (73°03.5´n 25°40.7´w). mountain in ne suess land between nanortalikdal and langgletscher. named during lauge koch’s 1947–49 expeditions by silvio eha for the shape (trug = trough). trumpington pas 72ø-515 (72°02.9´n 24°51.3´w; map 5). pass on the nw side of upper storgletscher, connecting with schuchert gletscher. named by the 1963 cambridge university expedition for trumpington street, cambridge, site of part of the university. (trumpington col.) trums ø 75ø-33 (75°58.4´n 20°10.3´w; maps 2, 4). island at the mouth of bessel fjord. so named by the 1906–08 danmark-eks pedi tionen, probably after the locality of the same name in nor way. (trums ö, trums island.) trumsdalen 76ø-163 (76°04.3´n 20°08.9´w; map 4). valley north of trums ø, where the bessel fjord hunting station (sometimes 329 called trumsdalen or trumsødalen) was established in 1932. the name was given by danish hunters who built a hut here in 1931. (troms dal.) træelv 73ø-180 (73°29.9´n 21°27.3´w). river on the south coast of hold with hope, named on an nsiu map (1932a; fig. 13) as treelva. fossilised wood was found in the delta of the river by norwegian hunters in 1929 (træ = tre = wood). wood valley has been used for the valley in which the river runs. trækpasset 76ø-42 (76°09.7´n 18°39.1´w; map 4). low lake-filled valley crossing store koldewey, named by the 1906–08 danmarkekspeditionen. the pass was discovered by hakon h. jarner in may 1907, and is one of the few places where it is possible to pull (= træk) sledges across the island. the name has also been said to have arisen from the wind through the pass (træk = draught). a hut built in the 1958 by danmarkshavn weather station at the east end of the pass (76°10.2´n 18°33.6´w) was reported to be in poor condition in 1971. (træk pass.) træsko sø 74ø (74°30.3′n 20°36.1′w). small lake in the area known as morænebakkerne, north of zackenberg forskningsstation. the name is used as a reference locality by scientists studying lake ecosystems. trønderheimen – see havna. tsavagattaq – see grundtvigskirken. tsulitsuuligai 69ø (69°54.6´n 22°56.2´w). name apparently used by greenlanders from scoresbysund for the over-wintering houses built on the south side of steward ø / sulussuutikajik in 1971–72; others have been built since, and there were four in 1993 (tuborg & sandell 1999). the name seems to be a modification of the official name for the island – see sulussuutikajik / steward ø. tuborgfondet land 78ø-30 (78°27.0´n 22°00.0´w; map 4). nuna tak in the garde nunatakker group west of nørre mellemland. named by the 1938–39 mørkefjord expedition after the committee of the tuborgfondet, which had made substantial donations to the expedition. (tuborgfondets land.) tugtut nunât – see tuttut nunaat. tunatinde 71ø (71°51.9´n 24°49.6´w; map 5). mountain on the south side of gannochy gletscher, central stauning alper. named by the 1968 university of dundee expedition which made the first ascent. tungen 80ø (80°31.8´n 19°46.0´w). glacier on the west side of the prinsesse caroline-mathilde alper, inner ingolf fjord, named by elmar drastrup’s 1938–39 expedition for its tongue-like shape. the name is used on the 1957 ams maps. tunnel pools 72ø (c. 72°13´n 24°00´w). name used by university of dundee expeditions between 1968 and 1974 for seven temporary lakes east of the lower end of tunnelelv gorge, west of mestersvig airfield. tunnelelv 72ø-196 (72°12.8´n 24°04.2´w; map 5). river draining store blydal, north scoresby land. so named by prospecting teams associated with lauge koch’s 1948–49 expedition, because it runs in a deep canyon at the mouth of the valley. falcons and geese regularly nest on the canyon walls. tunu 62ø-81ø official designation for east greenland in green landic. it appears to be a somewhat derogatory term as used by west greenlanders, and translates as ‘the back side’. tupikajik 70ø-150 (70°56.3´n 22°29.5´w). small hill on the valley floor near the south end of klitdal. one of the names recorded by the 1955 geodætisk institut name registration, the name refers to its shape, meaning ‘the little tent’. tupilaq 71ø (71°57.7´n 25°06.3´w; map 5). rock spire south of emmanuel fjeld in the upper reaches of sefström gletscher, stau ning alper. climbed by the 1998 scottish mountaineering club expedition, it was named after the inuit carvings traditionally made from sperm whale teeth. tuppiap qeqertaa [tobias ø] 79ø (79°20.6´n 15°46.5´w; maps 1, 4). island about 2 km long and 1500 m across with an ice cap about 35 m high situated in the greenland sea about 80 km from the coast of ne greenland. see also tobias ø. the island was named after the greenlander tobias gabrielsen. (tuppiat qeqertai.) turidsøen 73ø-587a (73°59.4´n 24°15.8´w). small lake in south ole rømer land, named by sigurd skaun and harald welde in 1932 as turidtjern. turner sund [immikkeertikajiip ikaasakajia] 69ø-21 (69°45.0´n 23°27.0´w). narrow sound separating turner ø on the northern blosseville kyst from the mainland. named by g.c. amdrup’s 1898–1900 expedition. turner ø [immikkeertikajik] 69ø-6 (69°42.0´n 23°24.0´w; map 3). island on the northern part of the blosseville kyst. named turner’s island by william scoresby jr. in 1822 in compliment and respect to dawson turner [1775–1858] of yarmouth, a wealthy banker, botanist and collector. (turner ö.) turnstone river 72ø (72°31.4´n 24°01.1´w). name used by the uni versity of dundee expeditions between 1968 and 1974 for a minor stream west of karupelv draining into holm bugt, sw traill ø. it was named for the birds (arenaria interpres). tusindstrinskløft 74ø-225 (74°04.6´n 21°35.9´w). minor ravine in nw hold with hope adjacent to river 7, on the north slope of frebold bjerg. so named by eigil nielsen during the 1931–34 treårsekspeditionen, presumably because one appeared to take a thousand steps to climb it. tuteinsfjæld 74ø (74°10.6´n 20°30.7´w). mountain on eastern clavering ø, on the south flank of rundetårn. the name appears on a sketch map in gustav thostrup’s 1921 logbook, and comme morates john tutein, a hunter of østgrønlandske fangstkompagni and an artist, who was killed by a bear while painting on 1 february 1921. he is buried at kap broer ruys. tutlas ø 76ø (76°37.8´n 20°37.7´w). skerry in west dove bugt, the present bratskæret. the name was proposed by the 1932 gefion expedition. (tutlas.) tuttut nunaat [renland] 70ø-27, 71ø-40 (71°15.0´n 27°00.0´w). land area bounded by nordvestfjord, øfjord, rypefjord and ed vard bay dal. the name was recorded by the 1955geodætisk institut name registration, and is a translation of the danish name, ‘reindeer land’. (tugtut nunât.) tuxensø 76ø (76°20.0´n 20°33.3´w). island in sw dove bugt, west of roon bugt, the present nanok ø. so named during the 1932 gefion expedition after henry tuxen [1890–1966], a director and civil engineer, who was one of the first committee members of na nok. repeated attempts by j.g. jennov to obtain approval of this, and several other nanok names, were rejected. the island was given the name nanok ø by the place name committee in 1940. tvegegletscher 74ø-178 (74°50.0´n 22°22.4´w; map 4). large gletscher west of th. thomsen land draining into svejstrup dal. the name derives from a sledge journey by th. johansen in early 1932, and was given because of the many tributary glaciers (tvege = fork). anna sten gletscher has also been used. tveholmen 73ø-252 (73°01.9´n 22°41.7´w). island in the broch øer group. so named on an nsiu map (1932a) because the island has two hills joined by a low narrow col. tvekegledal 71ø-80 (71°43.2´n 22°38.4´w). valley on the ne side of wegener halvø, so named during the 1931–34 treårsekspedi tionen by arne noe-nygaard because of the two cone-shaped (= kegle) mountains east of the valley. see also kegle i and kegle ii. tverådalen – see tværelv. tvillingbugt 74ø (74°09.3´n 20°22.2´w). twin bays separated by a small peninsula on se clavering ø, the east bay corresponding to lervig. the name appears on a sketch map in gustav thostrup’s 1921 logbook. tvillingegletscher 72ø (72°32.9´n 26°28.7´w). name occasionally used by haller (1955) for the twin glaciers østre tvillingegletscher and vestre tvillingegletscher south of the head of rhedin fjord. tvillingerne 70ø-216 (70°40.3´n 21°59.0´w; map 4). mountain 330 with twin peaks in south liverpool land. the name is said to have been given by aage nielsen during the 1924–25 expedition that found scoresbysund (storgaard 1926). it has also been attributed to captain vinther-jensen of one of the grønlands styrelse ships, and is reported as having been called by sailors as vinther jensen’s tvillinger. tvillingerne 76ø-17 (76°18.5´n 20°45.6´w). island in the sw part of dove bugt, so named by the 1906–08 danmark-ekspeditionen. it is divided into two parts by a low col. (tvillingøen, twin island.) tvillinghytten 76ø (76°19.0´n 20°48.3´w). norwegian hut built by john giæver’s expedition in august 1930 on the west side of tvil lingerne, sw dove bugt. it has also been known as kroken and nordlige jægersund-hytten tvillingnæs 77ø-86 (77°03.2´n 20°27.1´w). double cape on the north side of sælsøen. named by the 1938–39 mørkefjord expedition led by ebbe munck and eigil knuth. tvillingnæshytten 77ø (77°02.5´n 20°16.4´w). danish hunting hut on the ne side of sælsøen, built by nanok in october 1933 at tvillingnæs. now a ruin. it has also been known as sælsøhytten. tvillingodden 72ø (72°56.0´n 22°04.9´w). peninsula with two similar narrow projections on the north side of east geographical society ø. so named on the nsiu maps of lacmann (1937) (tvil ling = twin). tvillingøer 79ø (79°22.0´n 18°44.0´w). two islands off the ne coast of lambert land, the present eli knudsen øer. the name was used by the 1996 mylius erichsen’s mindeekspedition. dobbeltøer has also been used. tviskora 74ø (74°07.5´n 21°20.0´w). stream on south clavering ø west of eskimonæs. used on the nsiu maps of lacmann (1937), and so named because the stream has two (= tvi) outlets. tvivlsom 73ø (73°33.3´n 20°30.5´w). norwegian hunting hut on the east side of hold with hope, built in august 1927 by the foldvik expedition (tvilvsom = doubtful). it was also known as skandalen, bukta and moskusoksehytten. tværdal 70ø-188 (70°33.1´n 22°10.1´w). valley in south liverpool land draining south to hvalrosbugt. so named during the 1931–34 treårsekspeditionen by laurits bruhn, because it cuts deeply into the high plateau west of kronen (tvær = cross). tværdal 72ø-255 (72°56.5´n 23°04.1´w; map 4). valley crossing geographical society ø from sofia sund to vega sund, so named by desmond t. donovan during lauge koch’s 1949–50 expeditions. teigandalen has also been used. tværdal 73ø-50h (73°59.7´n 21°25.1´w). minor valley in river 14 on the north slope of stensiö plateau, nw hold with hope. so named by eigil nielsen during the 1931–34 treårsekspeditionen because it runs across the slope for some distance before joining river 14. tværdal 74ø-168 (74°20.9´n 20°33.1´w). valley on ne clavering ø, named by arne noe-nygaard and gunnar säve-söderbergh during the 1931–34 treårsekspeditionen in the form quertal (= transverse valley). tværdal 74ø (74°31´n 21°00´w). name occasionally used by danish hunters (drastrup 1932) for the present store sødal. tværdalen 74ø (74°11.4´n 20°17.3´w). valley running n–s across the east peninsula of clavering ø, connecting the valley containing henningselv with the valley running south into lervig. the name appears on a sketch map in gustav thostrup’s 1921 logbook. tværdalen 77ø-82 (77°17.3´n 21°16.3´w; map 4). valley extending westwards from annekssøen across okselandet. named by the 1938–39 mørkefjord expedition led by ebbe munck and eigil knuth, it was explored by paul gelting and alwin pedersen in june 1939. tværelv [aappaleqisaap kuua] 70ø-187 (70°33.1´n 22°10.1´w). river in south liverpool land draining tværdal. so named by lau rits bruhn during the 1931–34 treårsekspeditionen. tværelv 72ø-187 (72°11´n 24°10´w). river in north scoresby land, draining across the west side of blyklippen into store blydal. named by prospecting teams associated with lauge koch’s 1948– 49 expedition. tværelv 73ø-127 (73°48.3.0´n 20°51.0´w). river that forms the west boundary of home forland. derived from the tverådalen of the 1932a nsiu map, and so named because it cuts deeply into the plateau to form the lowest pass on the north side of tobias dal (tvær = cross). tværsund 71ø-112 (71°15.8´n 21°44.8´w). short sound on the sw side of trekanten, east liverpool land. the name was adopted from a suggestion by helge g. backlund during the 1931–34 tre års ekspeditionen. twelfth knight 71ø (71°08.6´n 26°36.5´w). summit 2055 m high on the ice cap between catalinadal and edward bailey gletscher, renland. climbed and named by the 2007 west lancashire moun t aineering group expedition. twincap 76ø (76°52.0´n 24°25.1´w). surveying station between two small ice caps in dronning louise land, the present shell iskappe and army iskappe. the name was used by the 1952–54 british north greenland expedition. tyrolerdal 74ø-320 (74°39.1´n 22°14.3´w; map 4). valley at the head of tyrolerfjord, between payer land and a.p. olsen land, named by louise boyd’s 1937 expedition. tyrolerfjord 74ø-65 (74°28´n 21°12´w; map 4). fjord on the nw side of clavering ø, extending inland north-westwards between payer land and a.p. olsen land. named by karl koldewey’s 1869–70 expedition as tiroler fjord, although in payer’s narrative (1876) the spelling tyroler fjord is used. the fjord was first explored by julius payer who named it for its magnificent alpine scenary, resembling the austrian tyrol. payer was a lieutenant in the austrian army, and was noted for his surveying in the austrian alps. the koldewey usage also included the present young sund, a name originating from william scoresby in 1822 that was reinstated by lauge koch in 1929. (tiroler fjords, tyrolerfjorden, tyroler fiord, tyrol fiord.) tyrolerfjord bundhytte – see bundhytten i tyrolerfjord. tyrolerheimen 74ø (74°21.8´n 21°51.7´w). original name of the norwegian hunting station at revet west of clavering ø, built in the summer of 1927 by the foldvik expedition (orvin 1930). a new station at the same locality, moskusheimen, was built by finn devold’s expedition in the summer of 1928. tyrolerheimen 74ø (74°25.8´n 21°26.9´w). a map in giæver (1930) suggests that the norwegian hunting hut on nw clavering ø usually known as bakkehaug may have gone under this name. this may be an error. tyrolerheimen 74ø (74°28.7´n 21°53.5´w). norwegian hunting hut on the south side of inner tyrolerfjord, built by the devold expedition in september 1928. the name is used by nsiu (1932c) and bang (1944). it is also known as skrænthytten. (tyroler-heimen.) tyskerdepot 76ø (76°40.6´n 18°43.8´w; fig. 17). name reported by fischer (1983) as used by staff at danmarkshavn for the remains of supplies put ashore by the 1944 goldschmied expedition (edel weiss ii) north of røseløbet, lille koldewey. danmarkshavn weather station made use of some of the fuel and coal in 1949 when the supply ship failed to arrive. the most conspicuous elements of the depot in 1990 were about 30 fuel drums, still bearing the clear insignia ‘kriegsmarine’ and ‘wehrmacht’. tyskit nunaat 71ø-230 (71°03.3´n 24°13.7´w). part of west jame son land south of gurreholm. recorded by the 1955 geodætisk institut name registration, the name means ‘land of the germans’. the eastern station of alfred wegener’s 1930–31 expedition was situated here, established with the assistance of greenlanders from scoresbysund. see also oststation. tyskit nunât – see tyskit nunaat. tyssa 73ø (73°31.4´n 20°38.8´w). minor tributary of the river glommen, se hold with hope. named on the 1932a nsiu map, 331 possibly for the place of the same name in the møre and romsdal district of norway, or derived from the dialect word for a gentle, singing noise, especially of a stream. tyvholmen 73ø (73°36.3´n 22°02.2´w). name occasionally used for the norwegian hunting hut in badlanddal built by the 1936–37 quest expedition, and more usually known as schelderup-hytten. tænderne 70ø-408 (70°21.2´n 25°30.7´w). mountain ridge west of kap stevenson with eight distinct summits resembling a row of teeth (= tænderne). named by w. stuart watt during the 1967–72 ggu scoresby sund expeditions. tärnholmen 73ø (73°28.0´n 21°30.9´w). small island or skerry in the inner part of mackenzie bugt, so named by kolthoff (1901) for the several hundred nesting terns he saw there in 1900. nsiu expeditions in 1932 used a similar name, tennholm, for the same reasons. the name also appears on some modern maps as tern holme or terneholm. (tärnön.) tærskeldal 72ø-429 (72°20.6´n 26°31.6´w). level, high valley between violingletscher and forsblad fjord. so named during the 1931–34 treårsekspeditionen by ove simonsen because the valley occurs at the watershed (tærskel = threshold). tærskelsø 72ø-483 (72°21.1´n 26°33.5´w; map 4). lake in tærskel dal, nathorst land. named during lauge koch’s 1954 expedition by john haller. tödibjerg 72ø-471 (72°07.6´n 26°44.3´w; map 4). mountain south of violingletscher, east of hjørnesø, nathorst land. named during lauge koch’s 1954–55 expeditions by hans zweifel after the moun tain tödi, the highest peak in the glarner alpen, switzer land. tøflerne 76ø-161 (76°44.0´n 19°02.9´w). two small skerries se of the trip-trap-træsko. discovered and so named during the 1932 gefion expedition, the name (= a pair of slippers) deriving from their proximity to the trip-trap-træsko (træsko = clogs). tölzer spids 71ø (71°51.0´n 25°16.5´w; map 5). mountain on the south side of the head of roslin gletscher. climbed by karl herligkoffer’s 1966 expedition on 15 august, and probably named after bad tölz, a small town in the bavarian alps, the home town of michl anderl, one of the climbers. (tölzer spids.) tømmerbugt 72ø-121 (72°52.5´n 25°09.1´w). bay on nw ella ø, sw of kap oswald. so named by the ella ø wintering party during the 1931–34 treårsekspeditionen because driftwood is common here. (tømmerbugten.) tømmerhuset 73ø (73°56.4´n 21°53.2´w). norwegian hunting hut built by the foldvik expedition in 1927 on the east side of loch fyne. it replaced the coffin-sized hut built by fritz øien known as villaen. tørelv 74ø-189 (74°18´n 21°50´w). river on west clavering ø. the name was used in the form dry river as a botanical reference locality by gelting (1934) during the 1931–34 treårsekspeditionen. tørvestakken 72ø-178 (72°55.7´n 23°00.0´w). one of the peaks of julekagen, geographical society ø. the name was given by the place name committee in 1939 for its resemblance to a pile of turfs. tørvedammen 74ø (74°28.7′n 20°33.2′w). locality in the vicinity of zackenberg forskningsstation. the name is used as a reference locality by visiting scientists. tørvekær 74ø (74°29.0´n 20°33.4´w). boggy area ne of zackenberg forskningsstation. the name is used as a reference locality by visiting scientists. (tørvekæret.) tågedal 72ø-145 (72°14.4´n 22°31.1´w). valley on extreme se traill ø on the north side of drømmebugten. so named during lauge koch’s 1936–38 expeditions by hans p. schaub because it was often filled by fog (= tåge). (taagedal.) tågefjeld 73ø-367 (73°45.0´n 24°33.8´w). mountain on the north side of the mouth of brogetdal, strindberg land. named by hans r. katz during lauge koch’s 1948–49 expeditions. (taagefjeld.) tågefjeldene 73ø-106 (73°41.0´n 21°14.0´w). mountain range in hold with hope. so named during the 1931–34 treårsekspedi tionen by th. johansen because the coastal fog banks often stop ped here and shrouded the tops (tåge = fog). tågefjeldene 80ø-47 (80°38.6´n 19°56.0´w; map 4). mountain range on the west side of inner ingolf fjord. so named by eigil nielsen during the 1938–39 mørkefjord expedition led by ebbe munck and eigil knuth, because they were often shrouded in fog (= tåge). (taagefjældene.) taagefjord 70ø (70°06´n 27°30´w). name used by hartz (1895) and gulløv (1991) for gåsefjord, because of the fog (= tåge) frequently encountered here during carl ryder’s 1891–92 expedition (tåge fjorden.) tågehjem 73ø-406 (73°19.8´n 25°39.7´w). plateau on the south side of benjamin dal, south andrée land. named during lauge koch’s 1948–50 expeditions by erdhardt fränkl for an occasion when his party was lost in the fog here. (taagehjem.) tågekyst 81ø (81°05.6´n 13°00.0´w). low coastal area of se kilen, kronprins christian land. the name is found on a coloured geo logical map of kilen printed in 1991 (pedersen 1991), and was suggested by christian hjort because it was often shrouded by coastal fog. tågeslugt 73ø-405 (73°18.7´n 25°39.9´w). ravine in south andrée land draining tågehjem. named during lauge koch’s 1948–50 expeditions by erdhardt fränkl. tågetoppene 73ø-698 (73°24.7´n 27°27.6´w; map 4). mountain summits up to 2240 m high in north frænkel land. so named during lauge koch’s 1949–51 expeditions by john haller, because they were often shrouded in morning fog (= tåge). tåkeheimen 74ø (74°59.0´n 18°23.7´w). norwegian hunting hut built in august 1952 by arktisk næringsdrift on the south side of shannon (tåke = tåge = fog). it was accidently burnt down in the 1980s. tårnet 71ø-393 (71°39.5´n 22°47.0´w). mountain on wegener halvø. named by katherina perch-nielsen during the 1967–72 ggu scoresby sund expeditions for its tower-like shape. eastern mountain has also been used. tårnet 72ø (72°07.1´n 24°58.7´w; map 5). peak about 2310 m high on the spiky ridge south of dansketinden, stauning alper. climbed and so named by the 1996 norwegian stauning alper expedition. the peak was climbed later the same summer by the 1996 scottish mountaineering club expedition who named it diannsketinden, although they measured an altitude stated as 2532 m. tårnfjeld 73ø-374 (73°44.0´n 26°30.5´w; map 4). mountain 2163 m high in andrée land, se of faustsøen. a spectacular rock wall with conspicuous white granites, the name arose during mapping by john haller and erdhardt fränkl during lauge koch’s 1948–50 expeditions. (taarnfjeld.) tårnfjeld 72ø-262 (72°14.3´n 24°37.8´w; map 4). mountain 2072 m high in the north stauning alper on the east side of skjold ungebræ. it was first climbed by a norwegian group from the west in 1951 and named taarnefjeld for the prominent rock tower near the summit. the name was subsequently adopted by john haller and approved at his suggestion. john hunt gave the name caerleon to the same peak, or an adjacent snow-capped summit, in 1960. the second ascent was made by an imperial college party in 1963. (tarnfjeld.) tårnfjeld gletscher 72ø-488 (72°12.6´n 24°35.7´w; map 5). gla cier on the north side of bersærkerbræ, north stauning alper. the name was apparently suggested by the place name committee as a substitute for john hunt’s 1961 proposal, caerleon glacier. tårngletscher 72ø-349 (72°09.8´n 22°31.1´w). glacier on se traill ø, so named during lauge koch’s 1956–58 expeditions by h.p. heres. 332 u udkiggen 73ø-104 (73°16.9´n 23°37.8´w; map 4). mountain in eastern gunnar andersson land, so named during the 1931–34 treårsekspeditionen by th. johansen because of the view. luncke fjellet has also been used. udkiggen 73ø-555 (73°00.6´n 27°46.6´w). mountain 2300 m high in goodenough land, named by j.m. wordie’s 1929 expedition as outlook peak. (udkiggen knolde.) udkigshøjen 76ø-60 (76°58.6´n 20°01.3´w). low hill near the se end of sælsøen, rising from a flat plain and providing a relatively good view. peter hansen had seen a musk-ox herd from the summit in the autumn of 1906. (udkigshöjen, udsigtshøj.) udkigspasset 73ø-561 (73°00.8´n 27°46.1´w). col on the north side of udkiggen, goodenough land. named by j.m. wordie’s 1929 expedition as outlook col, because after a period of bad weather a magnificent panorama was revealed. udsigtsryggen 73ø-331 (73°53.9´n 22°19.7´w). mountain ridge in east hudson land. so named by wolf maync and andreas vischer during lauge koch’s 1936–38 expeditions because of the view (= udsigt). ugla 73ø (73°31.7´n 20°53.1´w). small tributary of glommen in se hold with hope. so named on an nsiu map (1932a; fig. 13), possibly for places of the same name in the sogn & fjordane or sør-trøndelag districts of norway. ugleelv 70ø-113 (70°52.0´n 22°44.9´w; map 4). river nw of the head of hurry inlet. named by alfred rosenkrantz during lauge koch’s 1926–27 expeditions as owl river or ugle elv for the snowy owl (= ugle). ugleelv 74ø (74°29.0´n 20°30.8´w). reference locality ne of zackenberg forskningsstation. the name has been used by visiting scientists. uglehøjene 73ø-162 (73°31.1´n 20°50.2´w; map 4). mountain 429 m high in south hold with hope, named on an nsiu map in the form uglehaugane (nsiu 1932a; fig. 13), for the snowy owl (= ugle). uglespids 71ø-272 (71°56.0´n 23°50.5´w; map 5). mountain in the werner bjerge between the head of søndergletscher and østre gletscher, named by peter bearth and eduard wenk during lauge koch’s 1953–54 expeditions. it was climbed by bearth in 1953. the name was first published in the form eulenspitze (bearth 1954). uglesøen 73ø-661 (73°18.8´n 25°05.9´w). small lake on the south side of noa dal, ymer ø. the name originated during the 1931–34 treårsekspeditionen, and was approved at the suggestion of r. spärck. (uglesø.) ugpik-ravine 74ø (74°45.6´n 20°18.2´w). ravine in south kuhn ø on the west side of payer dal. named during lauge koch’s 1936–38 expeditions after the greenlandic word for a snowy owl. a pair had been found nesting on the wall of the ravine (maync 1947). ujaajiddudalaajik 70ø (70°27.5´n 21°58.0´w). name reported by the scoresbysund newspaper in 1984 as used locally for the large rock known officially as ujaatuk. ujaajiddudalajiip qammavia 70ø (70°27.9´n 21°56.6´w). name reported by the scoresbysund newspaper in 1984 as in use locally for fox havn, the bay on the east side of rosenvinge bugt officially known as ujuattuttalerajiip kangerterajiva. ujaattuttalerajik 70ø-316 (70°27.8´n 21°57.1´w). peninsula on the south coast of ujuattuttalerajiip kangerterajiva [fox havn]. one of the names recorded by the 1955 geodætisk institut name registration, it translates as ‘it has rather large rocks’. the scoresbysund newspaper recorded in 1984 the use of ujaajiddudalaajik. (ujái tug talerajik.) ujaatuk [store sten] 70ø-317 (70°27.5´n 21°58.0´w). large rock in the water on the east coast of rosenvinge bugt south of ujuat tuttalerajiip kangerterajiva [fox havn], southern liverpool land. recorded by the 1955 geodætisk institut name registration, the name means ‘the big rock’. (ujáituk.) ujáitugtalerajik – see ujaattuttalerajik. ujáitugtalerajîp kangerterajiva – see ujuaattuttalerajiip kanger terajivat. ujáituk – see ujaattuk. ujuaakajiip kangertiva [fønfjord] 70ø-18 (70°28.0´n 27°00.0´w). e–w-trending fjord between milne land and gåseland. one of the names recorded by the 1955 geodætisk institut name registration, it was named after the colony manager in scoresbysund, johan petersen, known as ujuât. the name translates roughly as ‘little johan’s fjord’. some modern maps record ujuaakajiip kangersua (tuborg & sandell 1999). (ujuâkajîp kangertiva.) ujuaakajiip nunaa [danmark ø] 70ø-67 (70°30.0´n 26°15.0´w). island at the mouth of fønfjord, off se milne land. the name was recorded by the 1955 geodætisk institut name registration, and is interpreted as ‘little johan’s land’. it was named after johan peter sen, colony manager in scoresbysund, known to the greenlanders as ujuât. (ujuâkajîp nunâ.) ujuaakajiip nunaata akia [gåsepynt] 70ø-69 (70°22.0´n 26°18.0´w). east cape of gåseland, south of danmark ø. one of the names recorded by the 1955 geodætisk institut name registration, it translates as ‘the point across from little johan’s land’. (ujuâkajîp nunâta akia.) ujuâkajîp kangertiva – see ujuaakajiip kangertiva. ujuâkajîp nunâ – see ujuaakajiip nunaa. ujuâkajîp nunâta akia – see ujuaakajiip nunaata akia. ujuattuttalerajiip kangerterajiva [fox havn] 70ø-314 (70°27.9´n 21°56.6´w). bay on the east side of rosenvinge bugt, south liverpool land. recorded by the 1955 geodætisk institut name registration, the name translates as ‘the little fjord with the large stones’. it is regarded as an unsuitable harbour because of the numerous rocks and skerries. the scoresbysund newspaper reported in 1984 the local usage as ujaajiddudalajiip qammavia. (ujái tug talerajîp kangerterajiva.) ukaleqarteq [kap høegh] 70ø-226 (70°43.4´n 21°33.3´w). peninsula on the east coast of south liverpool land. recorded by the 1955 geodætisk institut name registration, the name means ‘there are hares’. ukaleqarteq 70ø-205 (70°31.6´n 21°30.0´w). small valley or slope between kap lister and kap hodgson. recorded by the 1955 geo fig. 86. the island ulla ø at the mouth of grandjean fjord, looking eastwards. the john haller photograph collection, geus archive. grandjean fjord ulla ø 333 dætisk institut name registration, the name translates as ‘there are hares’. ukaleqartip oqqummut nuaa [snuden] 70ø-219 (70°41.0´n 21°34.9´w). cape on the east coast of south liverpool land, west of rathbone ø. one of the names recorded by the 1955 geodætisk institut name registration, the name means ‘the cape in the lea of ukaleqarteq’. (ukaleqartip orgungmut nûa.) ukaleqartip orgungmut nûa – see ukaleqartip oqqummut nuaa. ukattit kangersuat – see harefjord. ulddal 71ø-433 (71°11.2´n 28°54.0´w; map 4). valley in graben land west of eielson gletscher. so named by peter homewood during the 1967–72 ggu scoresby sund expeditions ggu be cause of the abundant musk-ox wool (= uld). musk oxen are present in large numbers on the nunatak of graben land. ulka 72ø (72°39.8´n 22°25.3´w). small island in east vega sund, east of nordenskiöld ø. used only on nsiu maps (lacmann 1937), and named after the ulke, a small fish. ulla ø 75ø-45 (75°06.7´n 20°53.3´w; map 4; fig. 86). island in the mouth of grandjean fjord, named after the wife of commander, later captain, asger emil valdemar grandjean [1889–1948]; she was hedwig alvine augusta (ulla) haak [1898–1986]. the name first appears on the 1932 edition of the 1:1 million scale geodætisk institut map based on lauge koch’s surveying. it has also been known as friedas ø. ullahytten 75ø-105 (75°07.6´n 21°03.3´w). danish hunting hut west of ulla ø, north of the mouth of grandjean fjord, built by nanok in september 1934. (ullestuen.) ullestuen – see ullahytten and olestua. ulmer spids 71ø (71°54.5´n 25°17.9´w; map 5). mountain about 2400 m high on the north side of uppermost duart gletscher, central stauning alper. first climbed by karl herligkoffer’s 1966 expedition on 17 august, and named after ulm in south germany, home of günter schweiger, one of the climbers. (ulmerspids.) ulstein 72ø (c. 72°52´n 21°50´w). small dark island 3 km se of kap mackenzie. the name is used in den grønlandske lods (1968), and is probably of norwegian derivation. see also ulsteinpynten. ulsteinpynten 73ø (73°55.3´n 20°14.7´w). west cape of jackson ø. the name appears on an nsiu map (1932a), and evidently derives from the village of ulstein, near ålesund in norway, the home town of peder sulebak who hunted in this region as a member of the hird expedition from 1929 to 1930. ulva 73ø (73°33.9´n 21°06.4´w). tributary of dyraelv in south hold with hope, so named on an nsiu map (1932a; fig. 13), for the wolf. ulvebjerg 80ø-78 (80°09.2´n 21°38.5´w; map 4; fig. 24). mountain in southern kronprins christian land, se of centrumsø. so named during lauge koch’s 1952–53 expeditions by erdhardt fränkl because fresh wolf (= ulve) tracks were found near the sum mit in august 1952. ulvebugthytten 75ø (75°01.6´n 21°28.1´w). name sometimes used for the danish hut built in september 1934 in central grandjean fjord which is officially known as grandjeanhytten. ulvedal 73ø-99 (73°32.9´n 22°19.6´w; map 4). valley on gauss halv ø at the west flank of giesecke bjerge. so named by th. jo hansen during the 1931–34 treårsekspeditionen because wolf tracks were repeatedly seen here (fig. 87). ulvedalen 71ø (71°12.2´n 23°11.1´w). name occasionally used on norwegian maps (ingstad 1935; akre 1957) for a valley in jameson land corresponding to the valley containing the present depotelv; it is the locality where helge ingstad observed tracks of two wolves (= ulve) in 1932. (wolf valley, ulvegjelet.) ulvedalen 72ø (72°52.4´n 25°05.6´w). name occasionally used for the ravine carrying kløftelv which drains ulvesø in nw ella ø. ulvedalen 74ø (74°12.5´n 20°23.5´w). valley on east clavering ø draining into grønnedal. so named on the nsiu maps of lacmann (1937) for the locality ‘ulvedalene’ near copenhagen. ulvedalene 76ø (c. 77°00´n 19°46´w). valley in germania land where wolves (= ulve) were seen. the name is used by poulsen (1991) in his diary of the 1906–08 danmark-ekspeditionen. ulvedræberhytten 73ø (73°39.3´n 20°52.7´w). danish hunting hut on reinaelv, hold with hope. it was built for nanok in 1939 (jen nov 1953) by christian petersen, also known as ‘ulvedræberen’ (= wolf-killer); he was also a fur trapper in canada for six years (p.s. mikkelsen 1994). ulveelv 70ø-128 (70°51.1´n 22°47.2´w). river west of the head of hurry inlet draing into ugleelv. it was named by alfred rosen fig. 87. adult wolf (ulv), a regular visitor to the centrumsø base camp in kronprins christian land in 1995. photo: jakob lautrup. 334 krantz during lauge koch’s 1926–27 expeditions in the form wolf river for a sighting of a wolf (= ulv) or wolf tracks. see also ulve dalen above. ulveheimen 74ø (74°21.8´n 21°51.7´w). name occasionally applied to the norwegian hunting station at revet, west of clavering ø at the head of rudi bugt. henry rudi poisoned a pack of eight wolves at revet in march 1930. ulvehøj 74ø (74°28.5´n 20°29.7´w). small hill 80 m high east of zac ken berg forskningsstation, and a name used as a reference locality by visiting scientists. ulvehøjen 75ø (c. 75°09´n 19°45´w). small hill, sometimes de scribed as having two summits, near the nanok hunting station in southern hochstetter forland. the name was used by danish hunters in the 1930s (e.g. nyholm-poulsen 1985). their foxtrap on the summit was reported to have caught principally falcons and owls. (ulvefjældet, ulvehøje.) ulvekam 73ø-707 (73°08.5´n 28°51.4´w). mountain ridge about 2200 m high west of petermann bjerg, in the nunataks of west frænkel land. so named during lauge koch’s 1951 expedition by john haller and eduard wenk because they encountered fresh wolf (= ulve) tracks here (fig. 87). (ulvebakken, wolfsrippe.) ulveodde 70ø-154 (70°51.1´n 22°27.9´w). small peninsula at the head of hurry inlet. named vargudden by a.g. nathorst’s 1899 expedition because two wolves (= ulve) were seen here on 5 au gust. n. hartz also saw two wolves here the following year (hartz 1902). (wolf point, vargodden, vargodde, wolf pynt.) ulveslugt 75ø (c. 75°19´n 17°50´w). one of the ravines at kap sussi where the 1943–44 operation bassgeiger excavated its subsurface base in a snow fan. the name is reported by olsen (1965), but as wolves were reportedly extinct in east greenland at this time it may not record a wolf sighting. ulvesø 72ø-126 (72°52.1´n 25°06.1´w). lake on nw ella ø, south of ella ø station. so named by the ella ø wintering party during the 1931–34 treårsekspeditionen because they found wolf (= ulve) tracks there. (wolf lake.) ulændedal 74ø-372 (74°38.5´n 22°33.4´w). valley on the sw side of pasterze, so named by the 1948 leeds university expedition because of its stony and rough character. it often contains an icedammed lake. (stony valley.) umimmakbjerg 71ø-72 (71°08.4´n 22°49.5´w). mountain in east jameson land. named during lauge koch’s 1926–27 expeditions by alfred rosenkrantz and tom harris as umimmak fjæld, a deri vation from the greenlandic for musk ox. (mt. umimmak fjæld, umingmakbjerg.) umimmalik 73ø-394 (73°34.4´n 23°08.4´w). valley in east andrée land draining south into the east end of grejsdalen. named during lauge koch’s 1948–50 expedition by erdhardt fränkl, it is green landic for musk-ox valley. (umingmalik) umingmak-ravine 74ø (74°46.0´n 20°17.2´w). ravine in south kuhn ø on the west side of payer dal. named during lauge koch’s 1936–38 expeditions for the greenlandic word for a musk ox, and used by maync (1947). umingmakbjerg – see umimmakbjerg. ûnarteq – see uunarteq. ûnarterajîp kangerterajiva, ûnarterajîp nûa – see uunarterajiip kang er terajiva, uunarajiip nuaa. ûnarterajik – see uunarterajik. ûnartertaqarteq – see uunartertaqarteq. ûnartertaqartikajîp orqunqmut kangertiva – see uunartertaqarti kajiip oqqummut kangertiva. ûnartip nûa – see uunartip nuaa. unicorn gletscher 77ø (77°11.0´n 24°00.0´w). name occasionally used for britannia gletscher in north dronning louise land. the name appears to have been that used during the 1952–54 british north greenland expedition, and is encountered in correspondance from expedition members. universitetets gletscher 72ø-522 (72°02.1´n 27°59.5 ´w). glacier in sw nathorst land draining south into leicester bugt. named by geoffrey halliday following botanical work during the 1961 leicester university expedition. see also leicester bugt. the name hammarskjöld bræ has also been used. the name is wrongly placed on some printed geodætisk institut maps. uomopasset 73ø-312 (73°49.4´n 22°08.4´w). pass in central hud son land between ankerbjergselv and ritomsø. named during lauge koch’s 1936–38 expeditions by heinrich bütler after passo del uomo, a pass near the ritomsee in switzerland. upper frederiksborg glacier – see øvre frederiksborg gletscher. uranus glacier 71ø-cl35 (71°32.8´n 25°16.7´w). name occasionally used in reports of james clarkson’s 1961 expedition and the 1962 oxford university expedition for the present oxford gletscher, south stauning alper. it is also used in some climbing reports (e.g. bennet 1972). many of the nearby glaciers are named after planets and stars, and the name is in keeping with these. ursus major gletscher 71ø-330 (71°41.5´n 25°20.5´w). glacier in the south stauning alper, a minor branch of jupiter gletscher. named ursa major glacier by john hunt’s 1960 expedition, after the constellation ursa major, or great bear. ursus minor gletscher 71ø-332 (71°40.0´n 25°15.1´w). glacier in the south stauning alper, a minor branch of jupiter gletscher. named ursa minor glacier by john hunt’s 1960 expedition, after the constellation ursa minor, or little bear. utbjørg 72ø (72°51.6´n 21°55.9´w). island off the coast of east geographical society ø, south of kap mackenzie. used only on nsiu maps (lacmann 1937), and named for its seaward position, and for a place of the same name in the vesterålen district of nor way. (utbjörg.) utburden 73ø (73°41.5´n 20°12´w). skerry 10 km off the east coast of hold with hope, so named on the 1932a nsiu map. the nor wegian word is employed for the ghost of an outcast or unbaptised child, or one born dead, a sinister conotation presumably inspired by the appearance of the skerry. utkiken 73ø (73°52.3´n 20°22.0´w). mountain 515 m high above kap james in wollaston forland. the name appears in this form on the 1932a nsiu map. it was used as a surveying station, and named for the view. uummataalaq [umimmakbjerg] 71ø-233 (71°08´n 22°49´w). mountain in east jameson land. the name was recorded by the 1955 geodætisk institut name registration, and translates as ‘the heart-shaped’. (umataulaq.) uunardaajik – see uunarterajik. uunardoq – see uunarteq. uunarteq [kap tobin] 70ø-323 (70°24.9´n 21°58.0´w). settle ment at kap tobin, south liverpool land. the name was recorded in 1933 by johan petersen, the first colony manager in scores bysund. it translates as ‘the place with hot springs’. the settlement took over some of the buildings of the radio station after it closed down in 1980, and in 1987 the population numbered 37. in 2007 there were no permanent residents. uunardoq has been recorded as the local spelling. (ûnarteq, ûnartok, onarteg.) uunarteq 70ø-325 (70°25.3´n 21°56.0´w). hot spring ne of kap tobin. recorded by the 1955 geodætisk institut name registration, the name means ‘the place with hot springs’. (ûnarteq.) uunarterajiip kangerterajiva 70ø-326 (70°25.5´n 21°55.2´w). bay east of kap tobin in south liverpool land with hot springs on both sides. one of the names recorded by the 1955 geodætisk institut name registration, the name means ‘the bay at the hot springs’. thala vig has also been used. (ûnarterajîp kangerterajiva.) uunarterajiip nuaa 70ø-328 (70°25.2´n 21°53.3´w). cape ne of kap tobin. recorded by the 1955 geodætisk institut name registration, the name means ‘the cape at the hot springs’. (ûnarterajîp nûa.) uunarterajik 70ø-327 (70°25.5´n 21°53.6´w). hot spring close to 335 kap tobin, south liverpool land. the name was recorded by the 1955 geodætisk institut name registration, and translates as ‘the not particularly hot spring’. the local population are reported to have used the variation uunardaajik. (ûnarterajik.) uunartertaqarteq 69ø-65 (69°35.0´n 23°42.0´w). se part of henry land, north blosseville kyst. the name was recorded by the 1955 geodætisk institut name registration, and translates as ‘the place with the hot spring’. (ûnartertaqarteq.) uunartertaqartikajiip oqqummut kangertiva 69ø-66 (69°34.0´n 24°10.0´w). fjord south of henry land, north blosseville kyst. one of the names recorded by the 1955 geodætisk institut name registration, it translates roughly as ‘the sheltered little fjord south of the place with the hot spring’. (ûnartertaqartikajîp orqunqmut kangertiva.) uunartip nuaa [kap tobin] 70ø-324 (70°24.6´n 21°56.7´w). southernmost cape of liverpool land. recorded by the 1955 geodætisk institut name registration, the name translates as ‘the cape at the hot springs’. (ûnartip nûa.) uunartoq qeqertaq – see warming island. 1v, 2v, 3v, 4v, 5v, 6v, 7v, 8v 72ø (72°07.2´n 23°55.8´w; map 5). designations used on 1:15 000 scale maps of the mesters vig region printed in 1951 for eight rivers west of expeditionshus flowing se into mesters vig. on some maps (e.g. bondam 1955) they appear as 1 vest – 7 vest. of these, river 2v was subsequently ap prov ed as holberg elv. v v. clausen fjord 77ø-33 (77°28´n 20°35´w; map 4). small fjord branch in the inner part of skærfjord, named by the 1906–08 danmark-ekspeditionen after viggo clausen [1875–1920], an of ficer in the danish navy (j. løve, personal communication 2009). (v. clausens fjord). vadrettal 72ø-440 (72°40.8´n 26°26.2´w). valley in gletscherland on the west side of rhedin fjord with a glacier at the head. so named by eugéne wegmann during the 1931–34 treårsekspedi tionen after a swiss locality of the same name. ‘vadret’ is a local italian/romansch name for a glacier. vagtpasset 73ø-560 (73°00.9´n 27°58.0´w). col between vedetten and knoen, goodenough land. named by j.m. wordie’s 1929 ex pedition as sentinel col because of its proximity to the mountain sentinel, now vedetten. vähfreude 72ø-459 (74°46.4´n 28°12.9´w). mountain west of fin delen sø, goodenough land. so named during the 1931–34 tre årsekspeditionen by eugène wegmann after ‘die käserei in der vehfreude’, a noted novel of the 19th century by jeremias gotthelf. vehfreude was a place very far away where the animals are full of joy. wegmann explored the area in august 1934. (vahfreude.) valborghytta 72ø (72°59.0´n 24°33.4´w). norwegian hunting hut in nw geographical society ø, built in september 1930 by arktisk næringsdrift. it is reported to have been named after one of the hunters’ homes in norway. it is also known as røvballehytten, svedenborg and joplassen. (valborghytten.) valdemarsmuren 77ø-108 (77°11.7´n 20°10.3´w; map 4). eastern edge of the highland area of søndermarken, forming the west border of slædelandet. named by the 1938–39 mørkeford expedition, the name derives from the valdemarsmuren, the main ram part of the dannevirke, sydslesvig, germany. the dannevirke earth walls date back to 800, while the main wall of stone was built by valdemar the great from 1160–80. valdemarshaab 74ø (74°15.9´n 19°23.0´w). danish hunting station built by østgrønlandske fangstkompagni at kap borlase warren in 1922, and taken down in 1923. it was replaced by a new hut built at sandodden, ny valdemarshaab, now known as sandodden. both stations were named after arner ludvig valdermar manniche [1867–1957], a director of østgrønlandske fangstkompagni. the station has also been known as kap borlase warren hytten and station. valhal 75ø-79 (75°01.0´n 22°23.4´w; map 4). ice plateau south of inner grandjean fjord. the name originated from the wintering party at kulhus during the 1931–34 treårsekspeditionen and was given for the ‘valhalla’ of nordic mythology, the home of the gods. valhallbreen 71ø (71°52.1´n 25°06.8´w; map 5). name given to a northern branch of roslin gletscher by the 1996 norwegian stau ning alper expedition. see valhal. valmuehytten 72ø (72°40.9´n 22°02.1´w). hut built by sirius in 1955 or 1956 on eastern geographical society ø, 3 km west of kap mc clintock (valmue = poppy). it is also known as kap mac clin tock hytten. van hauens fjeld 75ø (75°10.9´n 19°48.6´w). name occasionally used by danish hunters in the 1930s for nordre muschelbjerg, hoch stetter forland. james van hauen was a hunter with nanok from 1929 to 1931. vandfaldsklippe 71ø-435 (71°09.7´n 28°43.8´w). cliff on the east side of graben land, where a large river falls over the cliff edge to eielson gletscher. named by peter homewood during the 1967– 72 ggu scoresby sund expeditions. vandhulkløft 71ø-317 (71°38.9´n 24°37.3´w; map 5). ravine on the north side of the front of bjørnbo gletscher. so named by enrico kempter during lauge koch’s 1956–58 expeditions, be cause the river has eroded a hole through the permian arkoses to expose an inlier of crystalline rocks. vandreblok 74ø-309 (74°05.8´n 21°15.1´w). ice-transported boulder on the low ridge between østelv and østhavn, east of eskimo næs station, south clavering ø. the name originated from the wintering party at eskimonæs during the 1931–34 treårsekspedi tionen (vandre = wander, travel). vandreblokken 70ø-96 (c. 70°39´n 24°00´w; map 4). large, ice-transported boulder 15 × 10 × 5 m in size on the sw coast of jameson land. named by g.c. amdrup’s 1898–1900 expedition as vandreblok. vandredalen 80ø-60 (80°30.0´n 20°50.5´w; maps 1, 4). extensive n–s valley west of the alpine mountains of kronprins christian land, extending from marmorvigen to the head of ingolf fjord and northwards to romer sø. named by elmar drastrup’s 1938–39 expedition who traversed the valley in 1939 and considered it to be the likely migration route of musk ox between north and east greenland. it has also been viewed as an inuit migration route. vandrepasset 75ø-71 (75°55.8´n 21°58.0´w; map 4). pass in nørlund land between inner bessel fjord and knæksø. the name originated from the wintering party at kulhus during the 1931–34 treårsekspeditionen (vandre = wander, travel). vandskelsø 71ø-293 (71°51.3´n 26°54.3´w). ice-dammed lake in frederiksdal, nathorst land. named during lauge koch’s 1954– 55 expeditions by hans zweifel. the lake lies on the watershed (= vandskel), and when full can overflow southwards; when the glacier dam is broken it drains to the north. vandsø 76ø (76°46.5´n 18°42.6´w). name used by staff at dan markshavn for skibssø, which is the source of their drinking water. in spring and early winter aeroplanes bringing supplies and post sometimes landed on this lake. vandyke klipper 72ø-18 (72°07.6´n 22°20.6´w). the se part of traill ø was described by william scoresby jr. in 1822 as a stupendous cliff of singular beauty, with a prevailing colour of slate blue, intersected by zig-zag strata of bright yellow and red. he named it vandyke cliffs, probably because the colours and patterns reminded him of works by the notable dutch painter anthony van dyke [1599–1641], who is said to have altered the whole course of painting in england. (van dyk rock.) varde elv 70ø (70°36.1´n 22°37.4´w). name used by rosenkrantz (1934) for the river in vardekløft, on the west side of hurry inlet. varde nunatak 71ø-427 (71°11.0´n 29°16.4´w; map 4). nunatak 336 on the west side of vindue gletscher. so named by peter home wood during the 1967–72 ggu scoresby sund expeditions be cause of a cairn (= varde) on the summit. vardedalen 80ø (80°35.6´n 18°23.4´w). valley on the north side of ingolf fjord, ne of brede spærregletscher. so named by elmar drastrup’s 1938–39 expedition because they erected a cairn, scores by sundvarden, at the mouth of the valley. vardefjeld 72ø-261 (72°15.2´n 24°38.6´w; map 5). mountain in the north stauning alper on the east side of skjoldungebræ. it was first climbed by a norwegian party in 1951. the name was adopted by john haller following explorations during lauge koch’s 1954 expedition (varde = cairn). vardefjeld 73ø-156 (73°27.8´n 20°36.1´w; map 4). mountain 790 m high in se hold with hope. it appears on an nsiu map (1932a) in the form vardefjell, and was presumably named for a cairn. vardekløft 70ø-141 (70°36.1´n 22°37.4´w). ravine in neill klinter on the west side of hurry inlet. so named by g.c. amdrup’s 1898– 1900 expedition because the remains of a cairn (varde) built by carl ryder in 1891 were found on the cliff top above the ravine. (vardeklöft, vardekloeft, varde kløft.) vardekløft – see kløft i. vardenæs 74ø-49 (74°32.1´n 18°48.8´w). peninsula on the south side of sabine ø, east of germania havn. named by karl kolde wey’s 1869–70 expedition as cairn-spitze, because they built a cairn here with a report on the work of the expedition. (cairn spitze, cairn point.) varderyggen 76ø-232 (76°49.6´n 18°50.8´w; map 4). ridge in south germania land, nw of danmark havn. so named by the 1906–08 danmark-ekspeditionen, because a cairn was built here. (cairn ridge, varde-ridge.) vardepynten – see nuungajiva. vardevakt 72ø (72°25.0´n 24°33.8´w). norwegian hunting station 2 km se of kap peterséns built by the møre expedition in 1930 (rogne 1981). it is better known under the names sunnmøres heimen or kapp petersens. vargbukta 73ø (73°19.7´n 25°17.5´w). bay on the west coast of ymer ø, the present blomsterbugten. so named by nsiu in 1929, because the crew of the veslekari were surrounded by a pack of five wolves while capturing a musk-ox calf here. the name ap peared on several nsiu maps of the 1930s. (varg-bukta.) varghytta 73ø (73°19.9´n 25°16.9´w). norwegian hunting hut in blomsterbugten, west ymer ø, built by arktisk næringsdrift in march 1930. while building the hut olav kjelbotn and hallvard devold were surrounded by a pack of eight wolves, who subsequently followed them throughout their winter hunting trips. this was the same pack which henry rudi trapped with poison at revet later in the winter. see also vargbukta. the hut was maintained by sirius until 1979, and was restored by nanok in 2002. it has also been known as blomsterbugthytten. (wolf hut, vargheim, vargbukta.) vaskedalen 71ø-425 (71°16.8´n 29°03.0´w; map 4). valley west of vindue gletscher, where a large river flows from the glacier eastwards into an ice-dammed lake (vaske = wash). named by peter homewood during the 1967–72 ggu scoresby sund expeditions. vassdalen 73ø (73°36.0´n 22°38.7´w). valley west of ankerbjerg on the north side of moskusoksefjord, the present prospekt dal. it was used as a botanical reference locality in the report on nsiu investigations by vaage (1932). vassvika 72ø (72°55.7´n 22°07.5´w). bay on the north side of east geographical society ø, nw of kap mackenzie. the name is used only on nsiu maps (lacmann 1937), and was given for the boggy nature of the ground bordering the bay (vass = water). vastidal 73ø-81 (73°33.2´n 23°05.8´w; map 4). western of two exactly parallel valleys on central gauss halvø draining north to moskusoksefjord. named by lauge koch’s 1929–30 expeditions as vasti valley, or västidal, the name having been inspired by that of nearby gästisdal. the name means ‘western valley’. vauxhall 72ø (72°10.5´n 24°47.5´w; map 5). mountain 2140 m high between harlech gletscher and dunottar gletscher, north stau ning alper. first climbed by the 1963 imperial college expedition, and named after the london village, now swallowed up by lam beth, whose name is preserved in vauxhall bridge. vedel sø 76ø-332 (76°26.5´n 24°35.0´w; map 4; fig. 21). lake between pony gletscher and ejnar gletscher, dronning louise land. named by the 1952–54 british north greenland expedition after the danish vice-admiral aage helgersen vedel [1894–1981], head of the søværnet from 1950 to 1958 with special interests in the arctic and greenland. he was chairman of the dansk peary land ekspedition committee, which had assisted the british expedition during their preliminary 1951 expedition. vedet hytten 75ø (75°01.8´n 20°37.5´w). danish hunting hut about 2 km south of kap negri in fligely fjord, at the foot of the mountain vedetten. it was built by nanok in august 1951. it has sometimes been known as kap negri hytten. vedetten 73ø-551 (73°01.2´n 27°57.3´w). mountain 2200 m high in goodenough land, named by j.m. wordie’s 1929 expedition as sentinel for its appearance. vedetten 75ø-67 (75°03.0´n 20°41.2´w; map 4). mountain behind kap negri in ne th. thomsen land. the name originated from the wintering party at kulhus during the 1931–34 treårsekspedi tionen. vega sund 72ø-63 (72°53.5´n 24°00.0´w; maps 3, 4; fig. 12). sound between geographical society ø and traill ø. named by a.g. nathorst in 1899 as vegas sund, after the steamer vega, which carried the successful swedish expedition to spitsbergen in 1868, and through the ne passage in 1878–1880. (vega strait, vega sound, vegasund.) veganeset 72ø (72°50.1´n 23°10.0´w). peninsula on the north side of traill ø, equivalent to the present østernæs. so named by nsiu in 1929, because it lies on a pronounced bend of vega sund. the name has been used as a reference locality in danish botanical reports. the cape was later called kapp wollebæk in lacmann’s (1937) volume of norwegian maps. vegetation valley 77ø (77°32.6´n 20°47.9´w). valley in nordmarken, draining south to h.g. backlund fjord. named by the 1987 irish expedition to northern east greenland. vejle fjord 70ø-233 (70°45.5´n 21°42.0´w; map 4). fjord on the east coast of south liverpool land. so named during the 1931–34 treårsekspeditionen by laurits bruhn after the fjord of the same name on the east coast of jylland, denmark. vejrhøj 75ø-89 (75°03.2´n 22°54.7´w). nunatak sw of the head of grandjean fjord. the name derives from a sledge journey in 1932 by four men of the 1931–34 treårsekspeditionen along the margin of the inland ice. they were stranded by a storm here for two days. vendedalen 81ø-131 (80°59.2´n 20°00.0´w; map 4). valley on the west side of romer sø, kronprins christian land. so named by elmar drastrup’s 1938–39 expedition because the expedition turned back here in may 1939 (vende = turn). verenagletscher 72ø-293 (72°45.6´n 28°35.5´w; map 4). glacier in sw goodenough land. named by john haller following explorations during lauge koch’s 1953 expedition, it was said to commemorate a girlfriend of eugéne wegmann. (verena glet scher.) verbindungstal 74ø (74°42.2´n 22°13.6´w). valley connecting svej strup dal and tyrolerdal, the present mellemdal. the name was used by mittelholzer (1941) in his report on work during lauge koch’s 1938–39 expeditions (verbindung = connection). verena horn 72ø-463 (72°45.6´n 28°44.3´w). mountain in south goodenough land about 2320 m high. the name was used by eugéne wegmann during the 1931–34 treårsekspeditionen, and commemorates a locality in switzerland of similar name. it was climbed by john haller on 13 august 1953. (verena hornes.) vergys 73ø-322 (74°47.3´n 22°53.9´w). mountain in hudson land 337 south of dybendal, so named during lauge koch’s 1936–38 expeditions by heinrich bütler after the mountain chain of the same name in the savoy alps. verlorenes tal 72ø (72°27.0´n 22°00.0´w). name used by stauber (1938) for a valley on east traill ø, following work during the 1936–38 two-year expedition. it takes its name from a wild valley south of thusis, switzerland. in a slightly modified sense it was approved in the form ødedal, a name attributed to hans peter schaub (schaub 1942a, b). vermessungsbjerg 74ø (74°02.4´n 22°38.4´w). name used by helge g. backlund for the present rungstedbjerg, south of wordie bugt in north hudson land. vertebrae 72ø (72°07.7´n 25°09.8´w; map 5). small glacier on the north side of gully gletscher. probably named by the 1963 cam bridge university expedition, which climbed many peaks in this region. vesle finsch 73ø 74ø (74°00.3´n 21°07.0´w). next largest of the finsch øer, so named on an nsiu map (1932a) for its size relative to store finsch. (lille finsch ø, vesle finschøya). vesle vinteröya 73ø (73°11.5´n 23°00.0´w). smaller of the two vinter øer at the mouth of dusén fjord, so named on an nsiu map (1932a) for its relative size. vest kap 76ø (76°23.1´n 20°54.7´w). cape on the west side of gefion havn, godfred hansen ø. the name is used in den grøn landske lods (1968). vestelv 70ø-55 (70°40.2´n 25°32.6´w; map 4). river west of kap leslie, east milne land, named during the 1931–34 treårsekspedi tionen by hermann aldinger as westfluss. vestelven 74ø (74°28.6´n 20°36.3´w). reference locality used by vi si tors to zackenberg forskningsstation. vesterdalen 76ø-100 (76°46.8´n 18°48.6´w). valley west of danmarkshavn on the nw side of harefjeldet. so named by the 1906–08 danmark-ekspeditionen. (west valley.) vesterelv 74ø-255 (74°06.3´n 21°17.6´w). small river near eski monæs station, south clavering ø, draining into vesthavn. the name originated from the wintering party at eskimonæs during the 1931–34 treårsekspeditionen. flipa has also been used. vesterelv 80ø-64 (80°43.0´n 17°04.2´w; map 4). western of two rivers in south amdrup land. it was named informally as western river in the 1906–08 danmark-ekspeditionen reports, but was not approved until 1958. vesterelven 76ø-102 (76°46.3´n 18°41.4´w). western of the two rivers flowing into danmark havn near the original expedition house. named by the 1906–08 danmark-ekspeditionen as vesterelven. (vester elv.) vesternæsset 76ø-248 (76°03.3´n 20°05.2´w). peninsula about 7 km west of kap beurman on the north side of bessel fjord. it was used as a reference locality in the archaeological report of the 1906–08 danmark-ekspeditionen (thostrup 1911). (west naze.) vesterport 75ø-40 (75°18.3´n 21°15.1´w; map 4). mountain between femdalen and kildedalen in c.h. ostenfeld land. the name is attributed to the wintering party at kulhus in 1935. the mountain marks the gateway (= port) to the inner western part of the fjord, and the name apparently first appeared on a map in jennov (1939). kranges fjeld has been used for the same feature. vesterport sø 74ø (74°29.2´n 20°35.3´w). small lake in the area known as morænebakkerne, north of zackenberg forsknings station. the name is used as a reference locality by scientists studying lake ecosystems. vestersletten 73ø-190 (73°25.5´n 21°51.5´w; map 4). broad plain west of mackenzie bugt, named on an nsiu map (1932a) as vestflya. the 1932 nsiu expedition used the area as the main base for its flying operations, and established their balås flyveplass here. vestfjeldet 80ø-41 (80°35.0´n 21°27.7´w). mountain on the west side of sødalen, west of profilfjeldet. named by eigil nielsen during the 1938–39 mørkeford expedition as vestfjældet. vestfjord 70ø-14 (70°28.5´n 28°38.0´w; maps 3, 4). fjord extending westwards from the southern part of rødefjord. so named by carl ryder’s 1891–92 expedition for its direction. (west fjord, väst fjorden.) vestfjord gletscher 70ø-390 (70°18.0´n 29°24.0´w; maps 3, 4). large glacier at the head of vestfjord. named during lauge koch’s 1958 expedition by eduard wenk. vesthavn 74ø-257 (74°05.9´n 21°18.3´w). small harbour east of eskimonæs station, south clavering ø. the name originated from the wintering party at eskimonæs during the 1931–34 treårseks pedi tionen. (west harbour.) vestkronen 71ø-406 (71°52.9´n 23°39.9´w). mountain 1140 m high on the nw side of the bjergkronerne massif, north of ørsted dal. named by katherina perch-nielsen during the 1967–72 ggu scoresby sund expeditions. vestkæret 74ø (74°28.8´n 20°35.1´w). reference locality used by vi si tors to zackenberg forskningsstation. vestlandet 70ø (70°15.0´n 28°00.0´w). name used for the present gåseland in ragnvald knudsen’s diaries of carl ryder’s 1891–92 expedition to the scoresby sund region. vestmar bjerg 74ø-193 (74°13.3´n 21°18.3´w; map 4). mountain on sw clavering ø. the name was first used by gelting (1934) in the form mt. westmar, and was given for n.c. vestmar, captain of the gustav holm during the 1931–34 treårsekspeditionen. (vest mars bjerg.) vestplateau 74ø (74°02.1´n 21°39.5´w). minor plateau on the north slope of frebold bjerg, west of river 6, north hold with hope. so named during the 1931–34 treårsekspeditionen by eigil nielsen. vestporten 71ø-431 (71°13.2´n 27°55.9´w; map 4). mountain form ing the west side of edvard bay dal as seen from rypefjord. named by j.d. friderichsen during the 1967–72 ggu scoresby sund expeditions (port = gateway). vestre borggletscher 70ø-262 (70°05.9´n 23°50.4´w; map 4). glacier on volquaart boon kyst west of borgen, so named during the 1931–34 treårsekspeditionen by laurits bruhn. vestre brudelv 70ø (70°28.3´n 22°13.1´w). name used by alfred rosenkrantz for a west branch of brudelv, south liverpool land. vestre eskimovig 74ø-91a (74°05.7´n 21°11.5´w). bay west of eskimovig in south clavering ø. the name is used in the form west eskimo bay in the archaeology report of j.m. wordie’s 1926 expedition (johnson 1933). vestre gletscher 72ø-300 (72°01.7´n 24°08.0´w; map 5). western of three glaciers draining into the head of deltadal, north werner bjerge. the name first appeared on the maps of styger (1951) in the form vestregletscher, and stems from a climbing excursion during lauge koch’s 1950 expedition. vestre havnenæs 76ø-98 (76°45.4´n 18°42.5´w). peninsula on the west side of the mouth of danmark havn. so named by the 1906–08 danmark-ekspeditionen. (vr. havnenæs, west harbour pt.) vestre skanse 76ø-302 (76°57.6´n 20°05.9´w). plateau area west of pemmikanelv, south germania land. østre skanse occurs east of the river. named by the 1938–39 mørkeford expedition. vestre spærregletscher 72ø-450 (72°54.5´n 26°19.1´w). western of two glaciers in suess land which merge to dam murgangssø. adopted from a suggestion by eugéne wegmann who explored the region in 1933 during the 1931–34 treårsekspeditionen. vestre tvillingegletscher 72ø-285 (72°32.2´n 26°29.2´w; fig. 88). western of twin glaciers south of the head of rhedin fjord. named by john haller following explorations during lauge koch’s 1952– 53 expeditions. vestre vibeke gletscher 74ø-396 (74°20.8´n 24°17.9´w). western branch of vibeke gletscher, at the east side of batholin land. named by john haller following explorations during lauge koch’s 1956–58 expeditions. vestre vikingeborg 73ø-437 (73°03.9´n 26°43.6´w). mountain on 338 the west side of borggletscher, on the south side of kejser franz joseph fjord. named by john haller following explorations during lauge koch’s 1952–53 expeditions. vestreplateau 73ø-84 (71°04.5´n 21°41.4´w). small plateau 1000 m high west of margrethedal on gauss halvø. named by lauge koch’s 1929–30 expeditions as western plateau. vesttinden 74ø-192 (74°12´n 21°15´w). mountain peak on the nw side of taggletscher on sw clavering ø. the name was first used by gelting (1934) during the 1931–34 treårsekspeditionen, together with østtinden. vibeke dal 74ø (74°065.4´n 23°29.0´w). informal name used by sønderholm et al. (1989) for the valley in hudson land containing vibeke sø and vibeke elv. vibeke elv 74ø-329 (74°05.4´n 23°29.0´w; map 4). river draining vibeke sø, flowing eastwards through promenadedal to wordie glet scher. named by heinrich bütler during lauge koch’s 1936– 38 expeditions, originally in the form vibeckefluss. see also vibeke gletscher. (vibekes elv.) vibeke gletscher 74ø-301 (74°14.1´n 23°58.6´w; map 4). glacier between steno land and ole rømer land, dividing northwards into østre and vestre vibeke gletscher. mapped and named by lauge koch during flights in 1932 on the 1931–34 treårseks peditionen. vibeke nunatak 74ø-398 (74°22.1´n 24°11.5´w). nunatak be tween vestre vibeke gletscher and østre vibeke gletscher, east of bartholin land. named during lauge koch’s 1956–58 expeditions by john haller. vibeke sø 74ø-330 (74°08.5´n 23°46.0´w; map 4). large lake at the front of vibeke gletscher. named by heinrich bütler during lauge koch’s 1936–38 expeditions, and used first in the form vibekesee. (vibekes sø.) vibekefjæld 74ø (74°10´n 20°14´w). mountain on east clavering ø, part of magnetikerbjerg. the name appears on a sketch map in gustav thostrup’s 1921 logbook. girl’s name. victor madsen bjerg 73ø-82 (73°34.9´n 23°09.0´w). mountain on gauss halvø named during lauge koch’s 1929–30 expeditions as mt. victor madsen. see also victor madsen gletscher. victor madsen gletscher 73ø-588 (73°15.0´n 28°52.5´w; map 4). major n–s glacier between west frænkel land and martin knudsen nunatakker, which flows north, then swings east to join jættegletscher. mapped by lauge koch during flights in 1932 on the 1931–34 treårsekspeditionen, and named after victor madsen [1865–1947], director of the geological survey of denmark from 1913 to 1937. he was also on the committeee of the treårsekspedi tionen. (victor madsens gletscher.) vidarbreen 74ø (74°13.0´n 21°01.3´w). glacier on south clavering ø draining into skrællingedalen. used on the nsiu maps of lacmann (1937), the name is derived from old nordic mythology. vifteelv 70ø (70°28.6´n 22°11.3´w). name used by rosenkrantz (1942) for the small, fan-shaped river in south liverpool land draining gulfjelde (vifte = fan). vifteelv 72ø-523 (72°12.8´n 24°23.7´w). river flowing into skeldal on the east side of the stauning alper. the name was suggested by n.p. lasca following field work in 1966–67, and records the nume rous, large, depositional fans. vifteelv 77ø-85 (77°03.3´n 20°16.8´w; map 4). river on the north side of eastern sælsøen. named by the 1938–39 mørkefjord expedition, for the fan-shaped delta. vigdisdalen 74ø (74°18.0´n 21°40.0´w). valley on west clavering ø, draining into eigil elv and revet. used only on nsiu maps (lac mann 1937), the name is derived from an old norwegian personal name. vigfus dal 76ø-226 (76°57.8´n 21°45.5´w; map 4). valley at the head of mørkefjord. named by the 1938–39 mørkefjord expedition after vigfús sigurdsson [1875–1950], an icelandic farmer who looked after the horses used on j.p. koch’s 1912–13 expedition to the region. he also took part in alfred wegener’s 1930–31 eismitte expedition. vigfus elv 76ø (76°57.8´n 21°45.5´w). name used occasionally for the present mørkefjordselv. it occupies vigfus dal at the head of mørkefjord. vigfusdalfjord 76ø (76°57.0´n 21°27.6´w). name occasionally used fig. 88. two glaciers (østre tvillinge gletscher and vestre tvillinge gletscher) draining north-east from the lyell land ice cap into the valley between wahlenberg gletscher and rhedin fjord. the john haller photograph collection, geus archive. østre tvillingegletscher vestre tvillingegletscher 339 by the 1938–39 mørkefjord expedition for mørkefjord, into which vigfus dal drains. vikingebræ 72ø-96 (72°10.6´n 25°14.5´w; maps 4, 5). glacier in the north stauning alper draining west to alpefjord. named by ove simonsen during the 1931–34 treårsekspeditionen. vikingebugt [kangikajiip kangerterajiva] 70ø-75 (70°19.1´n 25°14.2´w; maps 3, 4). large bay between kap stevenson and helgenæs. so named by laurits bruhn during the 1931–34 treårs ekspeditionen. vildbjerg 73ø-442 (73°17.3´n 22°14.5´w). summit east of knuden and south of vilddalen, se gauss halvø. so named during lauge koch’s 1950 expedition by paul graeter (vild = wild). vildbæk 74ø-174 (74°13.7´n 21°34.1´w). small river on sw clav ering ø, a tributary to granatelv. the name appears to have first been used during the 1931–34 treårsekspeditionen, and was em ployed by malmquist (1932) in the form wildbach. vildbækdalen 74ø-263 (74°13.7´n 21°34.1´w). valley on sw clav e ring ø in which vildbæk flows. the name came into use during the 1931–34 treårsekspeditionen. kvisladalen has also been used. vilddalen 73ø-347 (73°17.9´n 22°13.1´w). valley in the south gie secke bjerge draining east. named during lauge koch’s 1936–38 expeditions by wolf maync and andreas vischer. skrukkedalen has also been used. (vildtal.) vildthorn 72ø-390 (72°02.3´n 23°15.9´w). mountain 1022 m high west of antarctic havn, north scoresby land. named by hans kapp during lauge koch’s 1957–58 expeditions, for its appearance. villa 71ø (71°57.0´n 22°44.1´w). original name for the norwegian hunting hut built in august 1930 by the møre expedition, nw of kap biot. it is also known as davy sund hytten and biot-stua. villaen 73ø (73°56.4´n 21°53.2´w). norwegian hunting hut on the east coast of loch fyne, north of strømmen, built by the foldvik expedition in 1927. the only pretentious thing about the original hut was said to be its name. it has been described as about the size of a large coffin, 2 m in length, 1 m high, and with a door in the roof (giæver 1958). it was built and used by fritz øien from 1926 to 1927, but demolished when a larger hut was built in 1927. the present hut has been known as norske villa, strømhytten and tøm mer huset. villaen 74ø (74°32.2´n 18°48.3´w). according to hvidberg (1932) this name was used for the house at germaniahavn, sabine ø, which he describes as the largest house in east greenland. it was also known as germaniahavn and blæsebælgen. villaen 76ø (76°46.2´n 18°41.1´w). name used by the 1906–08 dan mark-ekspeditionen for the expedition house built at danmark havn – see also danmarkshavnhuset. it was used by danish hun ters as a wintering station in the period 1919–21. the house is still standing, and now bears the name danmarksminde. vimmelskaftet 71ø-75 (71°42.7´n 22°44.0´w). minor valley on the nw side of wegener halvø, and also the name of the small house built at the mouth of the valley. the name appears to have been first used by spärck (1933) during the 1931–34 treårsekspedi tionen in the form vimmelskaft valley. the valley takes its name from the copenhagen street, which is narrow in its central part like a carpenter’s ‘vimmelbor’, a large drill with a long shaft. (vimmel skaftetdal.) vimmelskaftet 71ø (71°43.0´n 22°44.4´w). small wintering station built in 1931 by the 1931–34 treårsekspeditionen on the east side of fleming fjord, in the mouth of the valley vimmelskaftet. it is also known as kap brown huset and flemmingfjordhuset. vindblæsdal 70ø-392 (70°15.0´n 29°00.0´w; map 4). broad valley in the inner part of gåseland noted for the almost constant strong katabatic winds from the ice cap to the west. named by the 1963 geodætisk institut expedition. vindhjørne 70ø-374 (70°16.9´n 29°45.0´w; map 4). nunatak on the north side of vestfjord gletscher. so named during lauge koch’s 1958 expedition by eduard wenk, because of the strong winds experienced here. vindseløen 76ø-28 (76°48.5´n 20°19.3´w; map 4). island in nw dove bugt, so named by the 1906–08 danmark-ekspeditionen for its shape (vindsel = reel). (vindselöen, vindsel island, vindselseyra.) vindseløhytten 76ø-199 (c. 76°48´n 20°12´w). danish hunting hut on the east coast of vindseløen, nw dove bugt, built by nanok in november 1938. a very small hut, it has now disappeared. (vind selø hytten.) vindslugten 69ø-77 (69°48.0´n 26°24.0´w). outflow gap at the sw corner of geikie plateau, on the divide between magga bræ and bartholin bræ. it was named so by the 1969 watkins bjerge expedition, because they were held up here by a prolonged blizzard. (windy gap.) vindue gletscher 71ø-437 (71°14.0´n 28°55.0´w; maps 3, 4). glacier running along the west and north side of graben land. so named by peter homewood during the 1967–72 ggu scoresby sund expeditions because several geological ‘windows’ exposing older rocks occur beside the glacier. vindue nunatak 71ø-426 (71°14.7´n 29°06.2´w; map 4). nunatak west of vindue gletscher where older, geological formations are exposed by erosion through a thrust, a so-called ‘window’. named by peter homewood during the 1967–72 ggu scoresby sund expeditions. vindueskarmen 71ø-255 (71°55.7´n 23°38.0´w). mountain on the east side of the werner bjerge, south of blomsterdal. named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk, originally as sill-ryggen for the numerous basalt sills, subsequently corrupted to vindueskarmen (= window sill). vinkeldal 70ø-407 (70°39´n 26°20´w). broad valley on south milne land with several marked, right-angled bends. named by stuart watt during the 1967–72 ggu scoresby sund expeditions (vinkel = angle). vinkelklippe 77ø-131 (77°01.2´n 24°30.2´w; fig. 21). cliff on the north side of admiralty gletscher, dronning louise land, where the glacier makes a sharp turn to the ne. named by the 1952–54 british north greenland expedition as vinkel klippe. vinkelsø 72ø-431 (72°26.5´n 27°26.7´w; map 4). ice-dammed lake in the upper reaches of violingletscher, sw of cecilia nunatak. so named during the 1931–34 treårsekspeditionen by ove simonsen because of the right-angled shape of the lake (vinkel = angle). vintergata 74ø (74°16.9´n 20°56.9´w). large glacier on central clavering ø, the present skillegletscher. so named on nsiu maps of lacmann (1937) because when snow covered in the winter it can be used as a route to the interior of the island. vinterøer 73ø-29 (73°12.6´n 23°05.6´w; map 4). islands off the mouth of dusén fjord. named vinteröarne by a.g. nathorst in 1899 because they seemed a suitable place to overwinter, a good harbour and sheltered site for a hut being available. these islands were originally named broch inseln by koldewey 1869–70, but this name was moved by nathorst to islands off the mouth of sofia sund. (winter islands, vinter islands.) vinther jensen’s tvillinger – see tvillinger. violingletscher 72ø-424 (72°15.8´n 26°46.7´w; map 4). large glacier draining se into the head of furesø. named by eugéne wegmann who visited the area in 1933 during the 1931–34 treårs eks peditionen. the overall form of the glacier is s-shaped, as are groups of crevasses, reminiscent in shape of the head of a violin. vires acquirit eudo 70ø (70°47.1´n 22°08.3´w). minor peak 832 m high in liverpool land, west of bjerring pedersen gletscher. it was climbed by the 2002 loughborough grammar school expedition. the name is the school motto: ‘we gather strength as we go’. virgo gletscher 71ø-333 (71°41.2´n 25°11.0´w; map 5). small glacier in the south stauning alper, a minor branch of jupiter glet scher. named virgo glacier by john hunt’s 1960 expedition, after the constellation. 340 visdal 70ø (70°38.6´n 25°48.1´w). valley on se milne land draining into mudderbugt. the name, used on the maps of callomon & birkelund (1980), is said to derive from the strong forces which gave rise to the valley (vis = strength, in latin). visp 73ø-308 (73°55.8´n 23°39.5´w; map 4). river in west hudson land draining into johan davidsen dal. named by heinrich bütler during lauge koch’s 1936–38 expeditions after the river of the same name in the zermatt area of switzerland. vivian fjeld 76ø-44 (76°13.2´n 20°40.0´w; map 4). mountain 990 m high in northern ad. s. jensen land. named by henning bistrup during the 1906–08 danmark-ekspeditionen as vivians fjeld. pos sibly named after herbert vivian hertz, a colleague at the marine cadet school in 1898 (j. løve, personal communication 2009). (vivianbjærg.) vogt bjerg 74ø-239 (74°11.5´n 24°35.0´w; map 4). mountain in ole rømer land, named by sigurd skaun and harald welde during their 1932 expedition as j.h.l. vogt’s fjell, after one of wel de’s school teachers. vogt-hytta 72ø (72°37.0´n 22°38.4´w). hunting hut about 4 km west of kap palander on the ne coast of traill ø. built by arktisk næringsdrift in 1929. (vogtshytta, thorkild vogts hytta, thorolf vogts hytta.) volldal – see bjørnedal. volquaart boon kyst 70ø-74 (70°06.0´n 23°14.0´w; maps 3, 4). stretch of mountainous coast along the south side of scoresby sund between kap stevenson and kap brewster. the name was suggested by the geodetic institute in 1938 to commemorate the original discovery of scoresby sund by volquaart boon, a danish whaler aboard a dutch or german ship. when following the coast from 76°30´n to 68°40´n in 1761, the ship was swept by a strong current into a wide and deep fjord at about 70°40´n (bobé 1936). von krogh – see krogh-hytta. vrangelven 73ø-117 (73°55.8´n 23°58.9´w). river draining krum me langsø via johan davidsen dal to waltershausen glet scher. so named by sigurd skaun and harald welde during their 1932 expedition because they had great difficulty in crossing it (vrang = wrong, false). (vrangfluss.) vuachebjerg 73ø-318 (73°42.6´n 22°20.8´w). hill 400 m high on the west side of stordal in east hudson land. named by heinrich bütler during lauge koch’s 1936–38 expeditions, after the mountain vuacheberg near geneva, switzerland. (vuacheberg.) vuachehytten 73ø (73°41.6´n 22°09.7´w). danish hunting hut north of storelv, at the foot of vuachebjerg. it is also known as arvehytten and storelvhytten. vulkanhytten 73ø (73°45.5´n 20°59.6´w). danish hunting hut built for nanok in the spring of 1946 in the central part of tobias dal, hold with hope. possibly named for its proximity to a pingo, a glacial feature with conical shape (vulkan = volcano). it was originally known as jordly. vulkanhytten 74ø (74°00.0´n 22°12.1´w). danish hunting hut about 2 km inland from strømtangen, on the coast west of loch fyne. it was built by nanok in august 1951, and takes its name from the large conical pingo nearby. it is also known as dyndvulkan and jennovs næse. vædderen 76ø-25 (76°53.6´n 20°42.9´w; map 4). high peninsula between hellefjord and mørkefjord. named by the 1906–08 danmark-ekspeditionen. it was climbed by paul gelting on 25 july 1939 during the 1938–39 mørkeford expedition. the authorised spelling was changed from væderen to vædderen in 1973 to conform with its translation as ‘the ram’, and changed back to the original spelling with one ‘d’ in 1984. comprehensive official name lists from 1994 show a return to the ‘dd’ form. (the ram, vædder pla teauet, hrutfjall.) vædderhornet 76ø-220 (76°53.9´n 20°32.6´w). easternmost penin sula of vædderen. named by the 1938–39 mørkeford expedition, and used first by paul gelting on his journey in november 1938. it was approved in the form væderhornet up to 1973 and from 1984 to 1994. (vædderpynten.) vædderhytten 76ø-196 (76°51.3´n 20°44.6´w). danish hunting hut on the east coast of vædderen, north of the mouth of hellefjord. built by nanok in 1933, and rebuilt after an avalanche in 1938. the name was approved as væderhytten up to 1973 and from 1984 to 1994. væderen, væderhornet, væderhytten – see vædderen, vædderhornet, vædderhytten. vægtertårnet 70ø-443 (70°30.0´n 28°48.7´w). mountain 1340 m high between rolige bræ and vestfjord. named by laurent jemelin during the 1967–72 ggu scoresby sund expeditions for a resemblance to a watch-tower. vælddal 72ø-149 (72°19.2´n 23°02.1´w). valley on se traill ø, west of morris bjerg. named during lauge koch’s 1936–38 expeditions by hans peter schaub for a spring (= væld). (quelltal.) vælddalshytten 72ø (72°18.5´n 23°04.2´w). hut on the south coast of traill ø, on the east side of vælddal, built by nordisk mine selskab in the summer of 1970 from materials left at mestersvig by karl herligkoffer. vaagedalen 72ø (72°51.2´n 22°50.8´w). valley on south geograph ical society ø, a minor tributary to lysdal. so named on the nsiu maps of lacmann (1937) after jacob vaage [b. 1905], who participated as botanist on several nsiu expeditions, and was subsequently an editor with prominent interests in skiing. vågesund 77ø-21 (77°19.3´n 18°57.0´w). sound between rekved øen and the north coast of germania land. named by the 1906–08 danmark-ekspeditionen as vaagesund, possibly because a bear they shot here fell into a hole (= våge) in the ice, and was difficult to retrieve. w w. bishop sø 76ø (76°18.4°n 18°46.0´w). name used on 1952 wac maps for a lake on central store koldewey. w. horse-shoe mountain 71ø (71°39.2´n 22°21.2´w). part of the moun tain hesteskoen on canning land. the name is only found in the report by säve-söderbergh (1937). w. lynge skær 76ø-94 (76°43.6´n 18°29.4´w). small skerries se of danmark havn, named by the 1906–08 danmark-ekspeditionen after the archivist at søkortarkivet in copenhagen, denmark in the form w. lynges skær. wager nunatakker 69ø-78 (69°32.0´n 27°42.0´w; maps 3, 4). group of isolated nunataks south of scoresby sund, on the route followed by the 1969 watkins bjerge expedition. named after l.r. wager who had made some of the earliest geological mapping and climbs in the region. lawrence rickard wager [1904–65], a distinguished geologist, was professor at durham university from 1944 to 1950 and at oxford from 1950 to 1965. he is especially noted for his studies of the skærgaard intrusion. in 1935 he was one of the party that made the first ascent of the highest peak of the watkins bjerge (gunnbjørn fjeld / hvitserk; 69°55´n). wahlenberg gletscher 72ø-405 (72°30.0´n 27°00.0´w; map 4). glacier at the head of rhedin fjord. named by a.g. nathorst in 1899 as wahlenbergs glacier after georg (goran) wahlenberg [1780–1851], a swedish botanist, geographer and geologist, who was professor of medicine and botany at uppsala university from 1829. (wahlenberg glacier.) wallace bay 70ø (c. 69°57´n 22°25´w). the name was applied by wil liam scoresby jr. in 1822 to a pronounced bay on his chart sw of kap brewster, but as there are only indentations of the coast here of no great depth, the name has not been preserved. it was named after william wallace [1768–1843], who succeeded john leslie as professor of mathematics at edinburgh university in 1819. (wal lace bucht.) walter martin bjerg 71ø-345 (71°43.5´n 22°33.6´w). pyramid341 shaped mountain 608 m high sw of kap brown, wegener halvø. named by rudolf trümpy in memory of walter martin, a geology student from zürich who took part in the 1958 lauge koch expedition, and died in october 1959 in a climbing accident in the uri mountains. waltershausen gletscher 73ø-501 74ø-389a (74°00.0´n 24°40.0´w; maps 2, 4). major glacier 10 km wide between strindberg land and hudson land. so named by karl koldewey’s 1869–70 expedition, after baron wolfgang sartorius von waltershausen [1809–76], a noted german geologist who was professor of mineralogy and geology at the university of göttingen. (waltershausen-glacieren, waltershausen bræ, walters hausen glacier.) waltershausen nunatak 74ø-389 (74°15.0´n 26°15.0´w; map 4). large nunatak in the upper part of waltershausen gletscher. named during lauge koch’s 1956–58 expeditions by john haller. wapping 72ø (72°09.3´n 24°32.2´w; map 5). mountain 1680 m high on the east side of lower bersærkerbræ, north stauning alper. first climbed by the 1963 imperial college expedition, and named after the east london parish of wapping. warming island 71ø (71°28.9´n 21°51.5´w). island in northern liverpool land with three north-facing capes. it was formed by the melting of the ice cap to the south connecting it with liverpool land, and reported by dennis schmitt in 2005 as evidence for ra pid global warming. it has been given the unofficial name uunar toq qeqertaq in greenlandic. warming nunatak 74ø (74°24.2´n 23°29.9´w). name proposed during lauge koch’s 1929–30 expeditions for a nunatak in wordie gletscher already named faraway how by j.m. wordie. the name appeared on maps of seidenfaden (1931) and backlund (1932), and was given for johannes eugenius bülow warming [1841–1924]. he was a noted botanist, and professor at the university of copen hagen, denmark. washburns hus 72ø (72°13.3´n 24°03.2´w). name generally used by staff at mestersvig airfield for the house north of tunnelelv used by a.l. washburn as the headquarters for his geomorphological studies between 1955 and 1964 (washburn 1965). it has also been referred to as camp tahoe and det lille røde hus. watchtower glacier 71ø (71°12.4´n 26°25.1´w). glacier on the north side of edward bailey gletscher, renland. named by the 2007 west lancashire mountaineering group expedition. watkins bjerge 68ø-46 (69°00.0´n 29°30.0´w). mountain range lying almost entirely south of latitude 69°n, inland from the blos se ville kyst. this is one of the highest and most prominent mountain ranges in east greenland, rising to an altitude of almost 3700 m (gunnbjørn fjeld / hvitserk). the highest summit of the range has been convincingly argued by tornøe (1935) to be identical with the hvitserk of the icelandic sagas, although others (e.g. poul nør lund) had considered it improbable that hvitserk could be so far north. during their 1900 journey along the blosseville kyst and southwards to ammassalik, amdrup (1902b) reported seeing very high mountains looking northwards from the top of nordre aputiteq, with a pyramid-shaped peak that may have been the present-day gunnbjørn fjeld. in september 1930 the range was observed by gino watkins during a flight along the coast, and it was also observed on flights in 1933 by knud rasmussen and lauge koch. watkins originally called the range the new moun tains, while the name watkins land was used on a map compiled by lauge koch in 1933. the first ascent of gunnbjørn fjeld in the watkins bjerge was made by a party including ebbe munck, l.r. wager and a. courtauld in 1935. henry george (gino) watkins [1907–32] had attained an enviable reputation for his enthusiastic leadership of polar expeditions in labrador, spitsbergen and green land, notably the 1930–31 british arctic air route expedition (watkins 1932). he was drowned during his 1932 expedition to the ammassalik region of east greenland. watkins nunatakker 75ø (c. 75°45´n 22°45´w). land region west of ejnar mikkelsen gletscher, corresponding roughly to the present extent of kong wilhelm land. the name first appeared on the 1932 1:1 million scale geodætisk institut map prepared on the basis of aerial observations by lauge koch during the 1931–34 tre års ekspeditionen, and was given for gino watkins – see watkins bjerge. the name was dropped from later maps on the grounds that the region was not composed of nunataks, and the improbable grounds that there might be confusion with the wat kins bjerge south of scoresby sund. watson plateau 73ø-297 (73°35.2´n 23°30.0´w; map 4). plateau on west gauss halvø, named during the 1931–34 treårsekspedi tionen by gunnar säve-söderbergh after david meredith seares watson [1886–1973]. a british geologist who had described devonian fishes and early tetrapods, watson was for many years professor at university college, london. (watsons plateau.) weaselhytte 76ø-353 (76°39.7´n 19°40.7´w). hut on the south side of weaseløen, a small island south of søndre orienteringsø. the name was given by the staff at danmarkshavn weather station because the hut was transported to the site by the weasel tractors of the 1952–54 british north greenland expedition in march 1953. often damaged by bears, it was replaced in 1991 by a new hut, åndehullet. (weasel hut.) weaseløen 76ø-353 (76°40.0´n 19°40.8´w; map 4). small island in dove bugt. the name was reported by hans meltofte to be in general use by the staff at danmarkshavn weather station in 1969– 1971, because it was the site of weaselhytte. wedge 71ø (71°55.1´n 24°57.0´w; map 5). mountain 2340 m high on the ridge between storgletscher and dalmore glacier. named by the 1968 university of dundee expedition, who made the first ascent on 11 august. (wedge peak, the wedge.) wedge peak 71ø (71°38.5´n 25°17.9´w; map 5). rock peak about 2250 m high at the head of jupiter gletscher, west of tent peak, south stauning alper. first climbed by james clarkson’s 1961 expedition, and so named because of its bold outline, a stark crest of rock resembling a wedge glazed by slabs of ice. wefringdalen 72ø (72°57.8´n 24°25.0´w). valley on west geograph i cal society ø draining north into sofia sund. so named on the nsiu maps of lacmann (1937) after gunnar wefring [b. 1900], a norwegian artist who took part in several nsiu expeditions to svalbard and greenland. wegener halvø 71ø-87 (71°44.0´n 22°34.0´w; map 4). peninsula between fleming fjord and nathorst fjord. named by arne noe-nygaard during the 1931–34 treårsekspeditionen as wegener peninsula, after the german scientist alfred lothar wegener [1880–1930]. a german geophysicist and meteorologist, he took part in the 1906–08 danmark-ekspeditionen, the 1912–13 crossing of greenland led by j.p. koch, and died in 1930 on the inland ice during his own expedition. (wegener halbinsel, wegener halb insel, alfred wegeners halvö.) wegener øer 80ø-54 (80°34.3´n 16°46.6´w; map 4). group of small islands in the outer part of ingolf fjord. so named by eigil nielsen during the 1938–39 mørkeford expedition because alfred wegener built a cairn here during the 1906–08 danmark-ekspedi tionen. see also wegener halvø. wegenerflya 72ø (72°50.9´n 22°12.3´w). hillside on east geograph ical society ø, corresponding to the low-lying east slope of leitch bjerg. so named on the nsiu maps of lacmann (1937) after a.l. wegener. see also wegener halvø. weinschenck ø 77ø-141 78ø-50 (77°53.9´n 21°11.8´w; map 4). island nw of nordmarken. named during lauge koch’s 1956–58 expeditions by john haller, after a member of the 1906–08 dan mark-ekspeditionen. see also kap weinschenck. ggu’s new topographic maps show the island to lie entirely south of latitude 78°n. (weinschenk island.) weisse wand 72ø (72°03.4´n 25°06.1´w; map 5). name used by hans gesellman’s 1957 expedition for a mountain close to, or a 342 little se of korsspids, central stauning alper. the description of the first ascent in koglbauer (1965) reads as if a range of peaks covered by new snow was intended to bear the name. (weissen wand.) weisskopf 73ø (73°26.4´n 26°17.6´w). ice-capped mountain in southern andrée land. the name appears on a panorama drawn by john haller in 1949, reproduced in schwarzenbach (1993). it may have been intended as a tribute to john haller’s wife, susanne haller-weisskopf. wellenkamp spids 71ø (71°57.2´n 25°40.0´w; map 5). mountain on the west side of spærregletscher between castor glacier and pollux glacier. first climbed by the 1967 berchtesgadener expedition, who named it after j. wellenkamp, a mountaineer who made a number of notable climbs in the himalayas and andes in the 1950s. wendel pynt 76ø-99 (76°45.6´n 18°48.0´w). peninsula west of dan mark havn, south of harefjeldet. named by the 1906–08 danmark-ekspeditionen as wendels pynt, possibly after andreas c.d. wendel of the royal dockyard (j. løve, personal communication 2009). wenkhorn 73ø (73°25.1´n 26°14.5´w). mountain in southern andrée land. the name appears on a panorama drawn by john haller in 1949, reproduced in schwarzenbach (1993). it was in tended as a tribute to john haller’s professor at the university of basel, eduard wenk [1907–2001]. werenskioldflya 72ø (72°42.7´n 22°00.0´w). land area on se geographical society ø, immediately nw of kap mcclintock. so named on the nsiu maps of lacmann (1937) after the brothers werner and dagfin werenskiold. see dagfinvika and werner natnet. werner bjerge 71ø-197 (71°58.0´n 24°00.0´w; map 4). high mountain range in scoresby land east of the stauning alper. william scoresby noted these as the most elevated mountains he had seen upon the east greenland coast in 1822, and gave them the name werner mountains, in respect to the memory of the celebrated geologist, abraham gottlob werner [1750–1817]. the exact position of the mountains gave some difficulty to subsequent explorers, and they have been identified with some of the peaks of the stauning alper, and on some maps placed north of latitude 72°n. lauge koch fixed their approximate position in 1926–27 from a vantage point on the summits of traill ø. (werners moun tain, werners bjerge, wernerfjeld, werner bjærg.) wernervatnet 72ø (72°41.6´n 22°01.1´w). lake on se geographical society ø, wnw of kap mcclintock. used only on nsiu maps (lacmann 1937), the name was given for werner werenskiold [1883–1961], a norwegian geologist, geographer and glaciologist, and professor at oslo university, who participated in nsiu expeditions to svalbard from 1917 to 1924. (vernervatnet.) west gletscher 71ø (71°50.0´n 28°47.0´w). name used on 1957 ams maps for the present daugaard-jensen gletscher at the head of nord vestfjord. west icecap 69ø (69°30.0´w 26°10.0´w). name used in a report of the 1969 watkins bjerge expedition for the ice cap east of wager nunatakker, south of scoresby sund. west pond 72ø (72°14.4´n 23°55.0´w). name used by dundee uni versity expeditions between 1968 and 1974 for a small pool near langdyssen at the ne end of mestersvig airfield. westbrooks varde 76ø (76°35.2´n 19°02.8´w). cairn on the 760 m high summit south of the mouth of berg fjord, built in 1941 and with the nw side originally painted yellow. the name is used in den grønlandske lods (1968), and the cairn may have been built by the crew of the northland. the cairn was still standing in 1990. westendspids 74ø-29 (74°54.3´n 21°09.7´w). mountain 1404 m high at the west end of blåbærdalen. named by karl koldewey’s 1869–70 expedition as westend spitze, possibly because it was one of the westernmost points surveyed in the terrain mapped by the expedition. it was climbed by julius payer in 1869. (vest-spids.) western circus valley 73ø (73°09´n 23°14´w). name used by gunnar säve-söderbergh during the 1931–34 treårsekspedi tionen, together with eastern circus valley, for cirque valleys on the north slopes of celcius bjerg, ymer ø. western terrace – see western upper terrace. western upper terrace 73ø (c. 73°10´n 23°16´w). kochi-ridge, double ravine, eastern upper terrace, large debris cone, east plateau, kulisserna/coulisse and western upper terrace are a series of reference localities on the north and south slope of celcius bjerg, ymer ø. they were used during the 1931–34 treårsekspedi tionen by säve-söderbergh (1933). westernøya 72ø (72°44.2´n 21°55.6´w). island off the coast of se geographical society ø, north of kap mcclintock. used on the nsiu maps of lacmann (1937), the name was given for bjørn western [b. 1913], norwegian telegraphist on a 1932–34 hunting expedition to east greenland, and in 1935–36 telegraphist on jan mayen. (western øya.) westfal-larsen nunatak 73ø-575 (73°51.8´n 29°22.3´w; map 4). nunatak west of j.l. mowinckel land, named by arne høygaard and martin mehren in 1931 after the noted norwegian company founded by hans westfal-larsen in 1905. (westfal-larsens nuna tak.) westliche schwarze hügel 74ø (74°30.6´n 20°04.0´w). name used by vischer (1943) for the mountains on the west side of storsletten, wollaston forland. the name derives from work by wolf maync and andreas vischer during the 1936–38 two-year expedition. westminster 72ø (72°04.2´n 24°44.0´w; map 5). mountain 2500 m high between bersærkerbræ and schuchert gletscher, stauning alper, which bennet (1972) considered identical with royal peak climbed by the 1961 bangor jmc expedition. however, some climbers consider westminster to be a subsidiary summit a short distance east of royal peak. this was climbed and so named by the 1963 imperial college expedition, who gave it this name for the london district, since 1900 the city of westminster. weydmannsburg 72ø (72°03.4´n 25°06.1´w). mountain about 2700 m high on the ne side of sefström gletscher, so named and climbed by the 1964 aac zürich expedition. it is close to and may be the same as korsspids or weisse wand (bennet 1972). white 71ø (71°53.5´n 24°55.4´w; map 5). mountain about 2000 m high at the head of gannochy gletscher, central stauning alper. named for the colour by the 1968 university of dundee expedition who made the first ascent. whitefront pond 71ø (71°46.6´n 23°00.9´w). name used in an ornithology report of the 1963 british east greenland expedition (hall & waddingham 1966) for a lake on the north side of ørsted dal. a single greenland white-fronted goose was seen here on 18 july 1963. whittard bjerg 73ø-57 (73°49.4´n 22°36.1´w; map 4). mountain in east hudson land, named by lauge koch’s 1929–30 expeditions as whittard mtns after the chief geologist of james wordie’s 1929 expedition. the original usage was for a wider region including the present aravis and saussure massiv (seidenfaden 1931), but back lund (1932) restricted the name to the sw peak on the ridge. walter frederick whittard [1902–66] was professor of geology at bristol university from 1937, where he was noted for his encouragement of work in the arctic. (whittardfjellet, whittardberg.) wildspitze 75ø-15 (75°20.9´n 20°48.2´w; map 4). mountain 1599 m high in the southern barth bjerge. named during karl koldewey’s 1869–70 expedition, probably by julius payer, after the highest mountain in the otztal alps, austria. wildspitze was climbed in 1952 by members of the 1952–54 british north greenland expedition from a temporary base at kap rink, and in 1980 by members of exercise icy mountains vi, in both cases with the exception of the tottering 5 m summit tower. 343 wilkins nunatakker 74ø-175 (74°10.3´n 27°23.2´w; map 5). nunataks on the north side of eyvind fjeld gletscher. mapped by lauge koch during flights in 1932 on the 1931–34 treårsekspedi tionen, and named wilkins nunataks after george hubert wilkins [b. 1888], an australian pilot who with c.b. eielson made a pioneer flight with a lockheed vega in 1925 over the arctic ocean from barrow, alaska to green harbour, spitsbergen (wilkins 1928). william smith dal [adam af breemen dal] 72ø-257 (72°48.8´n 22°31.2´w). e–w-trending valley on geographical society ø between cambridge bugt and vega sund. so named by desmond t. donovan during lauge koch’s 1949–50 expeditions after william smith, the pioneer of stratigraphy known as the father of english geology, who was the first to make a geological map of england. the valley has another authorised name, adam af breemen dal, but this has rarely been used. brandal has been used for the same valley by norwegian scientists. wiman bjerg 73ø-112 (73°25.0´n 23°09.1´w). mountain on the south coast of gauss halvø. named during the 1931–34 treårs ekspedi tionen by gunnar säve-söderbergh as mt. wiman, after carl wiman [1867–1944], a swedish palaeontologist and stratigrapher. a professor at the university of uppsala, he was considered the initiator of swedish vertebrate palaeontology. norwegian maps of the 1930s used ramnefjeld for the same feature. windy corner 77ø (74°16.7´n 24°14.8´w). northern end of prins axel nunatak, where a party of the 1952–54 british north green land expedition was storm-bound for two days. it was known as a particularly windy area. the name was occasionally used informally in expedition accounts (simpson 1957). winge kyst 76ø-33 (76°50.0´n 19°15.0´w; map 4). sw coast of germania land, between snenæs and stormkap, a region where many of the detailed ornithological studies of the 1906–08 dan mark-ekspeditionen were carried out, and many features were named after birds. named after adolf herluf winge [1857–1923], a danish zoologist noted for his publications on greenland birds and animals, and who was vice-inspector of the zoological mu seum in copenhagen. winge assisted manniche (1910) in writing up his report. (winges kyst.) winston bjerg 76ø-312 (76°54.4´n 25°03.0´w; map 4; fig. 21). high mountain between admiralty gletscher and borg gletscher in west dronning louise land. the name was given by the 1952– 54 british north greenland expedition for winston churchill, who in 1952 was prime minister of great britain and a vice-patron of the expedition. he had made a substantial donation to the expedition. sir winston leonard spencer churchill [1874–1965], orator, author and statesman, is particularly remembered as the prime minister who led britain to victory in world war ii. wintherheimen 73ø (73°48.9´n 25°36.3´w). norwegian hunting hut built by levin winther in 1936 at the mouth of eremitdal, andrée land. levin winther [b. 1895], was a norwegian hunter who wintered in svalbard from 1928 to 1930, and in east greenland from 1935 to 1942, the last three years together with his wife petra. the hut has also been known as eremitdalhytten. wintherpasset 72ø (72°57.6´n 22°49.2´w). pass on central geo graph i cal society ø, in the northward extension of lysdal. so named on the nsiu maps of lacmann (1937) after levin winther. winthrop-young brae 72ø (72°06.2´n 27°26.9´w). small glacier at the west end of jomfrudal, west nathorst land. the name was used by geoffrey halliday during the 1961 leicester university expedition, and commemorates geoffrey winthrop young [1876–1958], a pioneer of british rock climbing. wittbergs bjerg 72ø (72°09.1´n 24°12.7´w). original name used for the mountain sw of mestersvig now known as schéele bjerg. it was named after carl ivar wittberg, rector of the engineering academy at filipstad, sweden from 1935 to 1957 (svend sølver, personal communication 2003). carl koch, who was responsible for con struction of the mine at mestersvig, was educated at the filipstad academy. as carl wittberg was still alive when the name was proposed to the place name committee, the name could not be approved, and schéele bjerg was substituted. wollaston forland 74ø-7 (74°25.0´n 19°40.0´w; maps 2, 4; fig. 15). large land area bounded by hochstetterbugten and young sund. named by william scoresby jr. in 1822 as wollaston foreland as a testimony of respect to william hyde wollaston [1766–1828], one of the commissioners of longitude. he was a chemist and physicist, and noted for his discovery of a process for making platinum mallable, which made him a fortune. (wollaston vorland, wollastone forland.) wood bjerg 71ø-27 (71°23.4´n 22°43.5´w). mountain 730 m high on the west side of carlsberg fjord. named by william scoresby jr. in 1822 as cape wood, after peter wood, a family friend and merchant with whom william scoresby sr. had business dealings. scoresby visited the woods at least once a week while he was at edinburgh university. wood valley 73ø (73°39.6´n 21°27.7´w). name used occasionally for the valley in which træelv flows in southern hold with hope. wordie bugt 74ø-275 (74°03.7´n 22°20.9´w; map 4). bay in the inner part of godthåb gulf, at the front of wordie gletscher. named in the form wordiebukta on the nsiu (1932a) map. james mann wordie [1889–1962] was a british polar explorer and petrologist who was chief of scientific staff on shackleton’s imperial trans-antarctic expedition 1914–17, visited spitsbergen in 1919, jan mayen in 1921 and east greenland in 1926 and 1929. in 1929 he made the first ascent of petermann bjerg. he was a founder member of the scott polar research institute, and its chairman from 1937 to 1955. (wordiebucht.) wordie gletscher 74ø-97 (74°15.0´n 23°05.0´w; maps 2, 4). large glacier draining into godthåb golf, named by lauge koch’s 1926–27 expeditions after j.m. wordie [1890–1962]. see also wordie bugt. (wordie glacier, wordies gletscher.) wordie kløft 73ø-50j 74ø-95 (73°59.5´n 21°22.9´w). ravine in north hold with hope, named by lauge koch’s 1926–27 expeditions as wordie creek after j.m. wordie, whose 1926 expedition had carried out important work here. see also wordie bugt. according to teichert & kummel (1976) koch’s original description gave no precise location for the ravine. rosenkrantz (1932) considered it to correspond to river 16, but nielsen (1935) to river 15. however, koch (1931) had called river 16 by the name blue river, now known in approved form as blåelv. wordie pas 72ø-498 (72°08.4´n 25°06.0´w; map 5). easy pass between the glaciers known as vertebrae and invertebrae, providing a link between gully gletscher and vikingebræ. named by the 1963 cambridge university expedition. see also wordie bugt. (wordie pass.) wordie’s cairn 73ø (73°07.5´n 27°14.3´w). cairn on the east side of the mouth of kjerulf fjord, built by j.m. wordie’s 1929 expedition to mark a fixed point in his survey of the region, and approximately on the site of one of the fixed points in dusén’s 1899 survey. the site is marked ‘cairn’ on wordie’s maps, and appears first in the form wordie’s cairn on the maps of louise boyd’s 1933 expedition (boyd 1935). the remains of the cairn were found by a ggu expedition in 1975 and rebuilt. see also wordie bugt. wordiesbugthytten 74ø (74°01.5´n 22°17.8´w). norwegian hunting hut built in 1936 on the south side of wordie bugt, about 2 km west of surprise elv in hudson land. it is also known as kalles hytte. (wordie bugt hytten.) wuss glacier 71ø (71°59.4´n 24°59.1´w). minor glacier on the west side of storgletscher, named by the 2007 smc east greenland expedition. 344 y yderbugten 76ø-95 (76°45.0´n 18°34.0´w). bay east of danmark havn, named in this form by the 1906–08 danmark-ekspedi tionen (yder = outer). hans meltoft reported in 1972 that the name østersøen was then in general use for this bay by the staff at danmarkshavn weather station. (yder bay.) yderdalen 73ø-120 (73°50.5´n 20°19.3´w). valley in home forland draining east to the coast south of kap james. named on the nsiu (1932a) map as ytterdalen, for its exposed position. yderhytten 76ø (76°35.7´n 18°44.7´w). danish hunting hut on the east coast of store koldewey, close to the low pass leading over to berg fjord. officially known as bergfjordhytten, it is also known as pashytten. the name yderhytten (= outer hut) is used to distinguish it from the nearby norwegian hut in berg fjord, also known as bergfjordhytten and inderhytten (= inner hut). yellow tor 75ø (72°25.4´n 20°59.5´w). mountain north of arden caple fjord, nørlund land, climbed by michael banks and richard brooke in 1952 during the british north greenland expedition (banks 1955). it was named for the yellow quartzites forming the summit. (yellow peak.) yllis 71ø (71°41.6´n 24°44.7´w; map 5). peak about 1881 m high in the south stauning alper between roslin gletscher and bjørnbo glet scher. climbed by the 1971 lancaster university expedition. ymer klinter 81ø (81°08.8´n 12°49.9´w). low cliffs in marine qua ternary sediments incised by the river anduin in east kilen, kron prins christian land. the name is found on a coloured geo logical map of kilen printed in 1991 (pedersen 1991), and was named after the ymer, the swedish ice-breaker that sailed along the coast in 1980. ymer nunatak 77ø-40 (77°24.8´n 24°16.1´w; maps 2, 4; fig. 21). large nunatak at the northern extremity of dronning louise land, named by the 1906–08 danmark-ekspeditionen as ymers nuna tak. ymer was a giant of norse mythology. ymer ø 73ø-26 (73°09.0´n 24°25.0´w; maps 3, 4). large island between sofia sund and kejser franz joseph fjord. named by a.g. nathorst in 1899 as ymers ø, after the swedish geographical journal ‘ymer’, which published many accounts of swedish expeditions to spitsbergen and greenland. see also ymer nunatak. (ymer is land, ymerøya, ymerinsel.) yngvar knudtzon fjeld 73ø-118 (73°56.1´n 23°48.8´w). moun tain in west hudson land, named by sigurd skaun and harald welde in 1932 as yngvar knudtzons fjell. young sund 74ø-9 (74°23.0´n 20°23.5´w; maps 2, 4). sound between clavering ø and wollaston forland joining up to the west with tyrolerfjord. it was originally named young’s bay by william scoresby jr. in 1822, in compliment to thomas young [1773– 1829], secretary of the board of longitude from 1818 to 1828. a physician and egyptologist, he was noted for his deciphering of hieroglyphics. karl koldewey’s 1869–70 expedition used tyroler fjord for the entire fjord, and young sund was reinstated for the outer part of the fjord by lauge koch about 1929 (seidenfaden 1931). (young’s bucht, youngsund, young inlet.) ytterhö 73ø (73°30.7´n 20°27.7´w). mountain 518 m high in se hold with hope, corresponding to the present rochusspids. so named on an nsiu map (1932a) because it is the easternmost and most exposed of this group of peaks. z zachariae isstrøm 78ø-13 (78°55.0´n 21°00.0´w; maps 1, 4). major glacier between hertugen af orléans land and lambert land. named by the 1906–08 danmark-ekspeditionen after georg hugh robert zachariae [1850–1937], a danish naval officer, later vice-admiral and director of the royal dockyard. he had also published a book on geodetic surveying (j. løve, personal communication 2009). (zachariaes isström, zachariaes bræ.) zachers grav 75ø (75°19.2´n 17°48.1´w; fig. 1943–44 bassgeiger). grave of gerhard zacher at kap sussi. he was a member of the german 1943–44 bassgeiger meteorological expedition, shot during a raid by the sledge patrol on 22 april 1944. the grave was intact in 1988, and the inscription on the broken cross still legible. a new, white-painted cross now marks the grave. zackenberg 74ø-63 (74°29.2´n 20°54.7´w; map 4). mountain 1338 m high on the north side of tyrolerfjord. named by karl kolde wey’s 1869–70 expedition, for its saw-tooth like summits. (mt. zachenberg, sachenberg, zachenbjergfjeldet, jagged mtn.) zackenberg 74ø (74°27.9´n 20°37.9´w). danish hunting station erected by nanok in zackenberg bugt in the summer of 1945. it was manned in the periods 1945–53 and 1959–60. the station is often used by sirius, and is in good condition. repairs were carried out by nanok in 1991–1992. it has also been known as horsnæs fangststation. zackenberg 70ø (70°44.5´n 22°51.7´w). name used for a mountain in jameson land north of j.p. koch fjeld by surlyk et al. (1973). zackenberg basen 74ø (74°27.9´n 20°38.4´w). house and store-hut built by the 1947–50 danish peary land expedition in zackenberg bugt immediately west of zackenberg hunting station. these facilities were also used by the 1952–54 british north greenland expedition, whose sunderland flying boats transported stores from here to britannia sø, dronning louise land. (basen.) zackenberg bugt 74ø-319 (74°27.5´n 20°38.9´w; map 4). bay on the north side of young sund, east of zackenberg, named by louise boyd’s 1937 expedition. zackenbergdalen 74ø (74°28.7´n 20°33.9´w). broad valley north of zackenberg bugt. the name is in common use by scientists visiting zackenberg forskningsstation. zackenberg forskningsstation 74ø-1000 (74°28.3´n 20°33.6´w). scientific field station north of zackenberg bugt adjacent to a gravel landing strip. the station was the concept of the danish polar center; it was built in 1995, officially opened in august 1997, and has since hosted visiting groups of scientists during the sum mer season. (zackenberg zero.) zackenberg-slette 74ø (c. 74°28´n 20°34´w). name used by various authors for the plain north of zackenberg bugt (e.g. christensen 1965; rosenberg et al. 1970). zackenbergelv 74ø (74°28.7´n 20°33.9´w). river draining store sødal, flowing east of zackenberg into zackenberg bugt. the name has been used by various authors in the past (e.g. jennov 1939), and has more recently come into regular use in the form zacken bergelven by scientists visiting nearby zackenberg forsk nings station. zackenberghytten 74ø (74°27.9´n 20°37.9´w). danish hunting hut in zackenberg bugt, se of zackenberg, built by nanok in july 1930 mainly for summer salmon fishing. zackenbergpasset 74ø (73°35.5´n 20°44.0´w). name used by wolf maync (1947) for the pass about 200 m high in lindemansdalen. wolf maync recorded the name as in use by danish and norwegian hunters during lauge koch’s 1936–38 expeditions. zackengrat 73ø (73°24.0´n 26°17.2´w). mountain ridge in southern andrée land. the name appears on a panorama drawn by john haller in 1949, reproduced in schwarzenbach (1993). zackengrat 74ø (74°21.5´n 20°37.4´w). basalt ridge about 560 m high near koralbjerg, east clavering ø. named during lauge koch’s 1936–38 expeditions by wolf maync and andreas vischer, and used in the report by maync (1942) where the name clearly refers to its jagged profile. zamiteselv 70ø-125 (70°53.4´n 22°43.8´w). river nw of the head of hurry inlet, named by alfred rosenkrantz during lauge koch’s 1926–27 expeditions as zamites river for the finds of fossil plants. zebra klint 71ø (71°51.3´n 24°00.9´w). prospector’s name used by nordisk mineselskab for a cliff adjacent to breithorn gletscher, 345 where quartz-barytes-galena veins with a distinctive striped intergrowth were found in 1971 (harpøth et al. 1986). zebra klippe 77ø-122 (77°13.1´n 24°49.3´w; maps 2, 4). northern cliff of iuel-brockdorff bjerg, dronning louise land. so named by the 1952–54 british north greenland expedition because it was formed of stripes of light sandstone and dark shale, which pro duced a distinctive zebra-like striped pattern. zechsteindal 73ø (72°25.1´n 22°06.4´w). name used by dunbar (1955) for the valley in the giesecke bjerge containing the river which wolf maync had referred to as zechsteinelv. zechsteinelv 73ø (72°25.1´n 22°06.4´w). name used by maync (1942) for a minor tributary on the north side of foldaelv, giesecke bjerge. it was named for the age of the rocks. zeissfjellet 74ø (74°19.3´n 21°04.3´w). mountain 1395 m high on central clavering ø, west of skillegletscher. so named on the nsiu maps of lacmann (1937), after the optical company carl zeiss, jena, germany, to commemorate their great advances in photo grammetric instrumentation. zeppelinfjellet 72ø (72°55.8´n 22°42.9´w). mountain ridge on cen tral geographical society ø, ne of lysdal, corresponding to the present langbjerg. used only on nsiu maps (lacmann 1937), the name was given for general graf ferdinand von zeppelin [1838–1917], who developed the airship for commercial services. zeus 71ø (71°41.3´n 25°08.8´w; map 5). rock peak 1850 m high on the south side of jupiter gletscher, south stauning alper. first climbed by james clarkson’s 1961 expedition and named after the chief deity of the greek pantheon. ziegla-husa 74ø (74°56.0´n 17°39.3´w). depot house built for the baldwin-ziegler expedition in 1901 at kap philip broke. it was inspected by the fiala-ziegler expedition in 1905, and ceded in 1930 to the norwegian state. with other norwegian hunting huts it passed to danish ownership in 1969. it has also been known as baldwin-huset and kap philip broke. the same name, ziegla-husa, has also been used for the similar huts built by the same expedition at bass rock (74°42.8´n 18°15.5´w). zielers sund 76ø (76°16.5´n 20°43.8´w). sound between tvil lingerne and the north coast of ad. s. jensen land, the present trangsund. the name was proposed by the 1932 gefion expedition, for kai zieler, a lawyer, and one of the committee members of nanok. zoologdalen 73ø-659 (73°21.0´n 24°17.7´w). n–s valley cutting across gunner andersson land, north ymer ø. the name originated during the 1931–34 treårsekspeditionen and was adopted at the suggestion of r. spärck. zuckerhütl 71ø (71°54.0n 25°40.3´w; map 5). mountain on the west side of spærregletscher, at the head of pollux glacier. first climbed by the 1967 berchtesgadener expedition, and named after the 3507 m peak of the same name in the ötztal region of the austrian tyrol. zurchergletscher 71ø (71°40.1´n 24°53.8´w). name used by stauber (1940) during lauge koch’s 1936–38 expeditions for a major glacier draining into schuchert dal, probably the present bjørnbo gletscher. zwergspids 71ø (71°50.7´n 25°23.1´w; map 5). small peak in the se corner of the upper basin of spærregletscher. climbed by karl herligkoffer’s expedition on 19 august 1966, and possibly named after the small town of zwergen in nw germany. (zwerg spids.) zwischenprofil 74ø (74°44.2´n 20°01.2´w). geological reference locality on se kuhn ø, used by maync (1947) in his description of work during lauge koch’s 1936–38 expeditions. æ æbeltoft vig 72ø-83 (72°30.0´n 22°10.0´w; map 4). bay on east traill ø, east of mols bjerge. named by ove simonsen during the 1931–34 treårsekspeditionen after the danish locality in the mols district, now spelt ebeltoft vig. ærenprisdal 71ø-391 (71°31.9´n 22°54.2´w). valley draining into pingel dal. the name was suggested by russel marris following his journey in 1968, and given for the common flowering plant of the figwort family. æselryggen 73ø-721 (73°55.8´n 22°20.5´w). ridge in ne hudson land, between stordal and loch fyne. the name is attributed to paul stern, who worked with lauge koch’s expeditions from 1955 to 1958 (æsel = donkey). æsken 71ø-399 (71°35.9´n 22°48.9´w). mountain 810 m high on south wegener halvø. named by katherina perch-nielsen during the 1967–72 ggu scoresby sund expeditions, for its angular shape (æske = box, case). ättestupan 73ø-523 (73°08.3´n 26°44.5´w; figs 68, 89). spectac ular near vertical cliff 1300 m high on the north side of inner kejser franz joseph fjord. this is the most impressive of the many high cliffs and capes in the central part of kejser franz joseph fjord. it was named by a.g. nathorst’s 1899 expedition after the cliff of the old viking legends, over which those who were tired of living or were afflicted by old age or illness cast themselves to death. (atte stupan, aettestupan.) ø 1ø, 2ø, 3ø, 4ø, 5ø, 6ø 72ø (72°08.6´n 23°47.4´w; map 5). designa tions used on 1:15 000 scale maps of the mesters vig region printed in 1951 for seven rivers east of expeditionshus which flow se into mesters vig. of these 1ø was subsequently approved as skibs elv. on other maps (e.g. bondam 1955), they were designated 1 øst – 7 øst. ättestupan fig. 89. the cliff ättestupan on the north side of kejser franz joseph fjord viewed from fjord level looking west. the cliff is about 1300 m high, and the fjord more than 1000 m deep. 346 øbjerg 71ø-276 (71°55.6´n 23°39.2´w). mountain in the werner bjerge on the south side of upper sirius gletscher, an island (= ø) in the ice. named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk. øbjerg 72ø-337 (72°28.6´n 22°08.1´w). mountain on se traill ø. so named by h.p. heres during lauge koch’s 1956–58 expeditions for its island-like position. øbugt 72ø (72°26.3´n 22°18.4´w). name used by stauber (1938) in a report on work during lauge koch’s 1936–38 expeditions for begtrup vig on the north side of mountnorris fjord. there is an island (ø) in the mouth of the bay. ødedal 72ø-169 (72°27.0´n 22°00.0´w). valley on east traill ø, between takkerne and kap parry, named during lauge koch’s 1936–38 expeditions by hans peter schaub for its barren appearance. (verlorenes tal.) ødedal 74ø-369 (74°55.3´n 21°41.9´w). valley in th. thomsen land on the south side of grandjean fjord. so named by the 1948 leeds university expedition for its desolate and bleak character, in contrast to grønningen to the south. (desolate valley.) ødegletscher 72ø-340 (72°27.1´n 22°04.4´w). small glacier draining into ødedal, se traill ø. named by h.p. heres during lauge koch’s 1956–58 expeditions. ødemarken 73ø-404 (73°22.1´n 25°54.0´w). high plateau in south andrée land on the north side of benjamin dal. so named by erdhart fränkl during lauge koch’s 1948–50 expeditions because of its desolate character, formed of limestone blocks that are very difficult to walk on. ødemarksdal 71ø-308 (71°31.1´n 24°47.0´w; map 5). valley west of karstryggen, west of schuchert flod, formed in desert-like, barren sandstone. named by enrico kempter during lauge koch’s 1956–58 expeditions. ødepas 72ø-138a (72°25.4´n 23°26.6´w). most prominent n–s pass in the svinhufvud bjerge, traill ø. named by lars b. clem mensen during a 1975–76 university of copenhagen expedition, to describe its barren and silent character. øen 76ø (76°46.1´n 18°40.2´w). area west of østerelv where it drains into danmark havn. it has also been called the island. øen 74ø-304 (74°05.2´n 21°16.7´w). small island very close to the south coast of the peninsula eskimonæs, south clavering ø. the name originated from the wintering party at eskimonæs during the 1931–34 treårsekspeditionen. øfjeld 73ø-386 (73°35.4´n 25°21.6´w; map 4). mountain in east andrée land north of grejsdalen. so named during lauge koch’s 1948–50 expeditions by erdhart fränkl because of its isolated island-like position. øfjord [ikaasakajik] 70ø-5 71ø-41 (71°00.0´n 26°12.0´w; maps 3, 4). long fjord between renland and milne land. discovered and named by carl ryder’s 1891–92 expedition during the exploration of the bjørneøer in september 1891 (fig. 7). there are no islands within the main stretch of the fjord, and the name derives from the bjørneøer group of islands at the ne end of the fjord. (öfjord, ö fjord.) öienfjellet 73ø (73°42.4´n 21°33.2´w). mountain 768 m high on the east side of loch fyne, equivalent to the present knasten. so named on an nsiu map (1932a) after fritz øien, a norwegian hunter who hunted in the loch fyne region for the 1926–28 foldvik expedition. he later spent five years on jan mayen as meteorologist and telegraphist during world war ii. øiens hus 73ø (73°40.6´n 21°44.9´w). norwegian hunting hut on the east side of southern loch fyne, built in august 1926 by the 1926–28 foldvik expedition. named after fritz øien, one of the hunters who helped build the hut. from about 1930 the hut was generally known as bunnhuset or botnhuset. øksebladet 76ø-39 (76°45.6´n 18°25.0´w). peninsula east of dan mark havn, so named by the 1906–08 danmark-ekspedi tionen because its shape resembles a two-sided axe. a hut between ørnen ø and øksebladet, built in 1949 by danmarkshavn weather station personnel, is sometimes known as øksebladet, and also as heering hus. (öksebladet, axe blade.) øksnevad 74ø (74°00.0´n 22°06.6´w). broad delta on the west side of the mouth of loch fyne. used only on nsiu maps (lacmann 1937), the name is derived from a norwegian dialect word for a place where cattle (in this case musk ox) go down to the water to drink. øresund 76ø-73 (76°42.1´n 18°39.1´w; map 4). sound between lille koldewey and kap bismarck. the name was used by trolle (1913) in his hydrographical reports of the 1906–08 danmarkekspeditionen, and was probably given for the sound of the same name between sweden and denmark. see also lille bælt and store bælt. (öresund.) øresundshytten 76ø (76°38.9´n 18°46.9´w). hut on the ne side of store koldewey, opposite røseløbet, built by the 1938–39 norsk– franske expedisjon. the name is misleading, as the sound it borders is lille bælt not øresund. it is also known as dagmar havn hytten. ørkenbjergene 71ø-79 (71°35.8´n 23°15.6´w). hills south of the head of fleming inlet, named during the 1931–34 treårsekspedi tionen by arne noe-nygaard as desert mts, because of their appearance. (desert bjergene). ørkendal 72ø-141 (72°56.2´n 25°21.9´w). valley in se suess land, in which the river kuukajik flows. named during the 1931–34 tre årsekspeditionen by ove simonsen, for its desert-like appearance. rud johansen valley has also been used. (örkedal.) orleans ø 77ø (77°43.0´n 17°45.0´w). name used for the present île de france (from 2004 qeqertaq prins henrik) by sophus poulsen during the 1906–08 danmark-ekspeditionen (lundbye 1984). the island was first mapped by the duke of orléans in 1905. ørnen ø 76ø-69 (76°44.0´n 18°26.8´w). island east of danmark havn, so named by the 1906–08 danmark-ekspeditionen, by chri stian b. thostrup after the danish navy petty officer association. (örnens ö, eagle island.) ørnereden 72ø (72°52.6´n 25°06.7´w). name often used for the main building of lauge koch’s ella ø scientific station built in 1931 during the 1931–34 treårsekspeditionen. the name translates as ‘eagle’s nest’. lauge koch was reputed to keep watch on the activities of expedition members with an eagle-eye from the main east-facing windows. örnereiret 73ø (73°58.7´n 21°17.4´w). norwegian hunting hut sw of the finsch øer, in hold with hope, built by the foldvik expedition in august 1926. the name appears in this form on the 1932a nsiu map, and translates as ‘eagle’s nest’. now said to be a ruin. (ørnereden.) ørsted dal 71ø-50 (71°47.5´n 23°12.0´w; map 4). broad, e–wtrending valley draining into fleming fjord. named by amdrup 1898–1900 as örsteds dal, after hans christian ørsted [1777– 1851], noted danish physicist and chemist. (örsted dal, örsted valley, ørsted valley.) ørsted dal hytten 71ø (71°45.6´n 23°23.8´w). norwegian hunting hut built by helge ingstad and normann andersen in ørsted dal, at the mouth of allday dal, in 1932–33. it was repaired in 1982 by otto lapstun, as a memorial to norwegian hunting activities. all day hytte has also been used. (ørstedsdal hytten.) østbræ 70ø-431 (70°10.0´n 25°58.4´w). minor glacier south of the mouth of gåsefjord, draining west to sydbræ. named during the 1967–72 ggu scoresby sund expeditions by e.a. hailwood. østerelv 74ø-254 (74°06.0´n 21°15.3´w). small river east of eski monæs station, south clavering ø. the name originated from the wintering party at eskimonæs during the 1931–34 treårs ekspedi tionen. trafsa has also been used. østerelv 80ø-65 (80°41.4´n 16°21.6´w; map 4). eastern of two rivers in south amdrup land. named originally as eastern river in a 1906–08 danmark-ekspeditionen report, and approved in the 347 present form in 1958. østerelven 76ø-101 (76°46.1´n 18°39.5´w). eastern of two rivers flowing into danmark havn near the original expedition house. named by the 1906–08 danmark-ekspeditionen as øster-elven. østernæs 72ø-76 (72°50.1´n 23°10.0´w; map 4). cape on the north side of central traill ø. the name appears to have been suggested by the place name committee in the 1930s as a substitute for the names kapp wollebæk and veganeset used by norwegians for the same feature. østerport 74ø (74°29.4´n 20°34.4´w). feature in the vicinity of zacken berg forskningsstation. the name is used as a reference locality in reports by visiting scientists. österreich-gletscher 72ø (72°00.4´n 25°03.7´w). name used in a report on hans gesellman’s 1957 expedition (koglbauer 1965) for the glacier on the ne side of sefström gletscher, more commonly referred to in mountaineering literature as kirkbrae. it was named after östereichspitze at the head of the glacier. österreichspitze 72ø (72°01.2´n 25°00.2´w). mountain about 2150 m high on the north side of sefström gletscher at the head of kirkbrae, a short distance north of bavariaspitze, stauning alper. named and first climbed by hans gesellman’s 1957 expedition. østersletten 73ø-107 (73°36.5´n 20°35.5´w; map 4). extensive, low-lying area in east hold with hope, named on an nsiu map (1932a) as aust-flya. østersøen 76ø (76°45.0´n 18°34.0´w). name at one time in general use by the staff at danmarkshavn weather station for the present yderbugten, east of danmarkshavn. hans meltofte reported that the name was used in correspondence, and by the catalina aircraft crew who used the locality as a landing site. østhavn 74ø-256 (74°05.6´n 21°16.0´w). bay east of eskimonæs station, south clavering ø. the name originated from the wintering party at eskimonæs during the 1931–34 treårsekspeditionen. eskimohamna has also been used. (east harbour.) østhytta 72ø (72°52.7´n 24°01.7´w). norwegian hunting hut on the south side of vega sund, north of rebild, built by arktisk nærings drift in 1929. the name arose about 1934 when it was the easternmost usable hut in vega sund. it has also been known as snøheim and traill hytten. now a ruin. (osthytta.) østkap 78ø-38 (78°42.8´n 19°09.3´w; maps 1, 4). southernmost of the several capes on the east side of schnauder ø, jøkelbugten. named by the 1938–39 mørkefjord expedition. østkap 76ø (76°24.8´n 20°45.0´w). name used on 1952 ams maps for the eastern cape of godfred hansen ø, east of the mountain sylen. østkronen 71ø-408 (71°49.3´n 23°24.7´w). mountain 1166 m high in the east part of the bjergkronerne massif, north of ørsted dal. named by katherina perch-nielsen during the 1967–72 ggu scoresby sund expeditions. østkæret 74ø (74°28.0´n 20°32.9´w). reference name used by visitors to zackenberg forskningsstation. östliche schwarze hügel 74ø (74°30.0´n 19°42.9´w). name used for the east side of storsletten, wollaston forland, by vischer (1943). it derives from work by andreas vischer and wolf maync during lauge koch’s 1936–38 expeditions. østporten 71ø-430 (71°11.7´n 27°44.2´w; map 4). mountain form ing the east side of edvard bay dal as seen from rypefjord (port = gateway). named by johan d. friderichsen during the 1967–72 ggu scoresby sund expeditions. østre borggletscher 70ø-265 (70°04.9´n 23°34.5´w). glacier on volquaart boon kyst east of borgen, so named during the 1931–34 treårsekspeditionen by laurits bruhn. østre brudelv 70ø (c. 70°28´n 22°12´w). name used by rosenkrantz (1942) for the east branch of brudelv, at the mouth of bruddal, south liverpool land. østre gletscher 72ø-302a (72°01.0´n 24°00.0´w; map 5). eastern of three glaciers draining into the head of deltadal, north werner bjerge. named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk. østre havnenæs 76ø-97 (76°45.3´n 18°39.3´w). peninsula on the east side of the mouth of danmark havn, so named by the 1906–08 danmark-ekspeditionen. (ör. havnenæs.) østre skanse 76ø-301 (76°57.0´n 20°03.6´w). plateau area east of pemmikankløft, south germania land. vestre skanse occurs west of the river. named by the 1938–39 mørkefjord expedition. østre spærregletscher 72ø-449 (72°53.7´n 26°07.7´w). eastern of two glaciers in suess land which merge and dam murgangssø. adapted from a suggestion by c. eugéne wegmann who explored the region in 1933 during the 1931–34 treårsekspeditionen. østre tvillingegletscher 72ø-286 (72°33.0´n 26°26.0´w; fig. 88). eastern of twin glaciers south of the head of rhedin fjord. named by john haller following explorations during lauge koch’s 1952– 53 expeditions. østre vibeke gletscher 74ø (74°20.6´n 24°05.4´w). branch of vibeke gletscher on the east side of vibeke nunatak. the name was given by john haller during lauge koch’s 1956–58 expeditions, but was not approved. østre vikingeborg 73ø-438 (73°03.8´n 26°36.3´w). mountain on the east side of borggletscher, on the south side of kejser franz joseph fjord. so named during lauge koch’s 1952–53 expeditions by john haller, for the castle-like appearance. østreplateau 73ø-88 (73°17.6´n 22°29.2´w). small plateau west of knuden and east of margrethedal on se gauss halvø, named by lauge koch’s 1929–30 expeditions in the form eastern plateau. østtinden 74ø-194 (74°09.2´n 21°08.8´w). mountain peak on south clavering ø. the name was first used, together with vest tinden, in a report by gelting (1934) on work during the 1931–34 treårsekspeditionen. østtungerne 77ø-57 (77°11.0´n 18°57.0´w; map 4). glacier in ne germania land on the east side of fladebugt. named by david malm quist during the 1931–34 treårsekspeditionen as öst tungerne. øtkerhytten 74ø (74°12.1´n 21°53.1´w). norwegian hunting hut on the sw coast of clavering ø, 5 km se of kap oetker, built by the foldvik expedition in 1927. it was replaced by a new hut in 1954. it has also been known as nes-odden and kap øtker hytten. øverbyefjellet 72ø (72°56.8´n 23°39.6´w). mountain 1150 m high on west geographical society ø. so named on the nsiu maps of lacmann (1937) after arne øverbye [b. 1906], norwegian tele graph ist on the 1931 and 1932 nsiu expeditions to east green land. (överbyefjellet.) øvre arkosedal 71ø-304 (71°34.5´n 24°45.3´w; map 5). upper part of the valley draining via nedre arkosedal to bjørnbo glet scher, with deep-red arkosic sandstone on both sides. named by enrico kempter during lauge koch’s 1956–58 expeditions. øvre frederiksborg gletscher 68ø-156 (69°00.0´n 31°32.0´w). name used for the glacier on the east side of frederiksborg nuna takker which extends to just north of latitude 69°n. the name was used by l.r. wager’s 1935–36 expedition in the form upper frederiksborg gletscher, as it is an upper northward extension of frederiksborg gletscher (wager 1937). the original name was given after the royal castle frederiksborg, hillerød, denmark. øvre gefionelv 72ø-185 (72°10.4´n 24°12.1´w; map 5). river in north scoresby land on the nw side of schéele bjerg, joining nedre gefionelv just before reaching store blydal. named by pros pecting teams associated with lauge koch’s 1948–49 expeditions. see also nedre gefionelv. (övre gefionelv.) øvre gefionpas 72ø-525 (72°10.2´n 24°15.2´w; map 5). pass be tween skeldal and øvre gefionelv. the name was suggested by n.p. lasca following work in the area in 1966–67. (gefion pass.) øvre mysteriesø 73ø-615 (73°15.3´n 28°11.0´w). higher of two lakes in mysteriedalen. louise boyd in 1933 distinguished wor die’s 1929 mystery lakes as upper mystery lake and lower mystery lake. øvre randgletscher 71ø-286 (71°52.7´n 24°07.4´w; map 5). upper and eastern of two glaciers south of aldebaran gletscher, on the north flank of randspids. named during lauge koch’s 1953– 54 expeditions by peter bearth and eduard wenk. øvre rypegletscher 73ø-547 (73°00.9´n 28°03.2´w). upper, netrending branch of rypegletscher, north goodenough land, named by j.m. wordie’s 1929 expedition as upper ptarmigan gla cier. øvre studer gletscher 72ø-307 (72°01.1´n 23°51.0´w; map 5). glacier in the north werner bjerge. named during lauge koch’s 1953–54 expeditions by peter bearth and eduard wenk. see nedre studer gletscher. øvresø 73ø-364 (73°46.7´n 24°55.1´w). small lake in brogetdal, strindberg land, above holmesø and laksesø. named during lauge koch’s 1948–49 expeditions by h.r. katz. øyedalen 71ø (71°52.7´n 22°57.6´w). broad valley in north scores by land south of antarctic havn, the present henrik møller dal. the name was used by norwegian hunters, and arose because the meandering river had left a series of ‘islands’ (= øyar). (øyadalen.) øyedalshytten 71ø (71°53.1´n 23°01.0). norwegian hunting hut built in 1932–33 for helge ingstad’s expedition in henrik møller dal, which norwegian hunters called øyedalen. öyneset 73ø (73°43.7´n 20°26.4´w). peninsula on the south side of carlshavn, eastern hold with hope, equivalent to the present knuds hoved. so named on an nsiu map (1932a), and possibly derived from a place name in the aust-agdar district of norway. å aage bertelsen gletscher 80ø-114 (80°17.0´n 19°35.5´w; fig. 24). glacier on the north side of hekla sund. named by john haller following explorations during lauge koch’s 1956–58 expeditions after aage bertelsen [1873–1945] – see also kap aage bertelsen. aage de lemos dal 72ø-92 (72°46.1´n 24°06.9´w; map 4). valley on nw ymer ø. the name was suggested by ove simonsen in 1983, and given for aage de lemos. a long-serving member of lauge koch’s geological expeditions, he was telegraphist from 1931 to 1942, station leader on ella ø 1933–42 and equipment chief 1947– 59. he was a member of the sledge patrol in 1941–42, and was sometimes referred to as the ‘king of ella ø’. de lemos wintered at the station on ella ø for a longer period than anyone else, made many climbs around ella ø, and surveyed the skerries in vega sund. aage nielsen fjæld 70ø (70°30.3´n 22°10.1´w). name used by rosen krantz (1934, 1942) for one of the summits of gulfjelde in south liverpool land. see also aage nielsen gletscher. (mt. aage nielsen.) aage nielsen gletscher [apusiikajik] 70ø-215 (70°40.2´n 21°48.9´w). glacier in se liverpool land. so named by laurits bruhn during the 1931–34 treårsekspeditionen after aage nielsen [1902–26], a young astronomer who overwintered at scoresbysund during the expedition to found the colony in 1924–25, and died soon after returning to denmark. (åge nielsen gletscher.) aagenæsfjellet 74ø (74°21.0´n 20°47.6´w). mountain on north clav ering ø, equivalent to the present koralbjerg. the name is used on the nsiu maps of lacmann (1937), and was given for sigurd aagenæs [1905–33], a norwegian pilot who took part in the nsiu expedition in 1932. åkerblom ø 72ø-33 (72°29.3´n 24°37.8´w; map 4). island at the mouth of segelsällskapet fjord, named by a.g. nathorst in 1899 as åkerbloms ö after filip åkerblom [1869–1942]. he was a geophysicist, subsequently professor of meteorology at the university of upp sala from 1907 to 1934. åkerblom acted as meteorologist, hydro grapher and physicist on the 1899 expedition. (aakerbloms ö, aker bloms island, åkerblomöya.) ålborg fjord 71ø-99 (71°38.5´n 22°08.5´w). fjord or large bay in east canning land, named during the 1931–34 treårsekspe di tionen by arne noe-nygaard as aalborg fjord after the town in jylland, denmark. ålborghus 76ø-205 (76°23.3´n 20°54.4´w; map 4). danish hunting station at gefion havn on the south side of godfred hansen ø. built in august 1938 by nanok with funds raised by the danish newspaper ‘aalborg stiftstidende’, and named originally in the form aalborghus. it replaced a hut on the same site built in 1933. the station was manned in the periods 1938–41 and 1945–52, and was maintained by sirius until 1988. (aalborghus station.) åndehullet 76ø (76°39.7´n 19°40.8´w). hut built by staff of dan markshavn weather station on the south side of weaselø, as a re placement for weaselhytten. aanstadpasset 72ø (72°55.7´n 23°35.5´w). pass on western geo graphical society ø, so named on nsiu maps of lacmann (1937) after the norwegian botanist sigurd aanstad [b. 1906], who took part in the 1932 nsiu expedition to east greenland. århus bugt 71ø-98 (71°44.0´n 22°06.0´w; fig. 90). bay or fjord in kap tyrrell århus bugt kap wardlaw canning land nathorst fjord fig. 90. view southwards of the eastern peninsulas of canning land, kap tyrell and kap wardlaw, separated by the ice-covered waters of århus bugt. the john haller photograph collection, geus archive. 348 349 north canning land, named during the 1931–34 treårsekspedi tionen by arne noe-nygaard as aarhus bugt after the town in jyl land, denmark. noe-nygaard was a student at aarhus kathedralskole. aarsethsundet 72ø (72°42.9´n 22°44.7´w; fig. 14). sound between silja ø and south geographical society ø, in vega sund. used only on nsiu maps (lacmann 1937), the name was given for elling aarseth [b. 1897], a norwegian ship-owner who supplied ships used by nsiu expeditions. åsen 75ø-68 (75°14.4´n 19°46.9´w). hill in southern hochstetter forland. the name originated from the wintering party at kulhus during the 1931–34 treårsekspeditionen. (åsen = the ridge). aasesøen 73ø-585 (73°59.5´n 24°22.1´w). lake in south ole rø mer land, named by sigurd skaun and harald welde in 1932 as åsevannet. girl’s name. (aasesee.) 350 abbreviations ams: army map service, corps of engineers, us army, washington, d.c., usa awi: alfred wegener institute for polar and marine research, bremer haven, germany bgr: bundesanstalt für geowissenschaften und rohstoffe, federal institute for geosciences and natural resources, hannover, ge r many cedme: centre for studies and documentation on polar areas, dijon, france dmu: danmarks miljøundersøgelser, danish environmental re search institute ecopolaris: successor to the cedme organisation from 2003, based in dijon, france gbu: grønlands botaniske undersøgelse, greenland botanical sur vey geus: de nationale geologiske undersøgelser for danmark og grønland, geological survey of denmark and greenland gfm: grønlands fiskeriog miljøundersøgelser, greenland envi ron mental research institute grea: groupe de recherches en écologie arctique, arctic ecology research group ggu: grønlands geologiske undersøgelse, geological survey of green land gi: geodætisk institut, geodetic institute gto: grønlands tekniske organisation, greenland technical orga nisation icao: international civil aviation organization kms: kort & matrikelstyrelsen, national survey and cadastre nanok: østgrønlandsk fangstkompagni nanok a/s, east greenland trapping company nanok ltd. nsiu: norges svalbardog ishavsundersøkelser, norwegian sval bardand arctic ocean survey sirius: slædepatruljen sirius, sirius dog sledge patrol usaf, wac: united states air force, world aeronautical charts, aeronautical chart and information center, st. louis, usa vildtbiologisk station, kalø: wild game station, kalø zackenberg zero: zackenberg ecological research operations, zackenberg, greenland geographical terms bjerg, bjerge: mountain, mountains bræ: glacier bugt: bay dal, dalen: valley, the valley elv: river, stream fangsthytte/station: hunting hut/station fjeld, fjelde, fjeldet: mountain, mountains, the mountain gletscher: glacier halvø: peninsula havn: harbour hus, huset: house, the house hytte: hut, cabin kap: cape klint: cliff klippe: crag, cliff kyst: coast land: land næs: headland, cape pynt: point skær: skerry spids: pointed summit strand: beach sund: sound sø: lake tinde: pinnacle, peak ø, øer: island, islands glossary administrative organisation of greenland grønlands styrelse (statsministeriet) – greenland administration under the ministry for state (1925–50) grønlandsdepartement (statsministeriet) – greenland department under the ministry for state (1950–55) ministeriet for grønland – ministry for greenland (1955–87) hjemmestyre – home rule (1979–2009) selvstyre – self-government (2009–) aarseth, s. 1993: staunings alper. american alpine journal 1993, 168–169. akre, b. 1957: fri manns liv, 144 pp. oslo: gyldendal norsk forlag. akre, b. 1983: ‘heltene’ i kald krig, 109 pp. oslo: eget forlag. aldinger, h. 1935: geologische beobachtungen im oberen jura des scoresbysundes (ostgrönland). meddelelser om grønland 99(1), 128 pp. alendal, e. 1980: overføringen af moskusfe fra øst-grønland i perioden 1899–1969. polarboken 1979–80, 73–92. amdrup, g. 1902a: beretning om skibsexpeditionen til grønlands østkyst fra d. 14. juni til d. 18. juli og fra d. 12. september til d. 4. oktober 1900. meddelelser om grønland 27(2), 183–271. amdrup, g. 1902b: beretning om kystexpeditionen langs grønlands østkyst 1900. meddelelser om grønland 27(4), 109–152. amdrup, g. 1913: report on the danmark expedition to the northeast coast of greenland 1906–08. meddelelser om grønland 41(1), 270 pp. andersen, f.s. 1937: ueber die metamorphose der ceratopogoniden und chironomiden nordost-grönlands. meddelelser om grøn land 116(1), 95 pp. andersen, j. 1980: kaptajn ejnar mikkelsens mindeekspedition 1980. grønland 1980, 295–315. andersen, j. 2005: grønland – på eventyr i kajak, 299 pp. københavn: gyldendalske boghandel, nordisk forlag a/s. andersen, t. 1960: gensyn med scoresbysund. grønland 1960, 353–359. andreasen, c. 2003: ringen sluttet. in: martens, g. et al. (eds): peary land, 172–178. nuuk: forlaget atuagkat. anrick, c.-j. 1932: en sommarresa til östgrönland. ymer 52, 175–212. anonymous 1985: staunings alper. american alpine journal 1985, 220 only. anonymous 2000: ofjord region. tsavagattaq, south ridge and other activity. american alpine journal 2000, 241 only. arke, p. 2003: scoresbysundhistorier. fotografier, kolonisering og kort lægning, 164 pp. valby: borgens forlag. 351 references bmc report archive: british mountaineering council, freepost nat 11244, manchester m20 7za, uk. a british organisation working for climbers, hill walkers and mountaineers, bmc receives reports from expeditions that it has advised or supported. since about 2000, these reports have been available online and can be downloaded in pdf format (www://thebmc.co.uk). dpc report archive: polarbiblioteket (the polar library), strand gade 102, dk-1401 copenhagen k, denmark. following the closure of the danish polar center (dpc) in 2009, many staff and some functions were transferred to the ministry of science, technology and innovation, but after widespread protests the dpc library facility (now known as polarbiblioteket) was preserved on the ground floor of the original dpc building. the library holds a large col lec tion of unpublished expedition reports dating from about 1973 to 2008, most of which were submitted during the period when dpc was officially responsible for granting permission for scientific and sporting expeditions to greenland. submission of a re port after the return of an expedition was one of the conditions of being granted a permit. geus archive: geological survey of denmark and greenland (geus), øster voldgade 10, dk-1350 copenhagen k, denmark. the sur vey holds a large collection of material arising from the activities of staff geologists and summer contract geologists in greenland and denmark, including those of the former geological survey of greenland (ggu) and the former geological survey of denmark (dgu). the survey also holds several hundred black and white prints of photographs taken from norseman aircraft by lauge koch’s geological expeditions in the 1950s; the photograph collection was formerly held by john haller, chief geologist of lauge koch’s expeditions, and was donated to the survey by john haller’s widow. a much larger collection of several thousand negatives of aerial photographs taken during lauge koch’s expeditions is held by the geological museum, copenhagen. kort & matrikelstyrelsen (kms: national survey and cadastre): kms incorporates the former geodætisk institut (geodetic insti tute), and continues to have responsibility for production of maps of greenland. however, after the transfer of the archives of the place name committee to greenland (see below) it has retained only a short run of the place name committee minutes (1967–80). there is very little documentation of the work of the place name committee at the danish rigsarkivet (state archive), apart from standard names lists and other documents widely distributed by the former geodetic institute. place name committee archive: considerable documentation of the work of the place name committee for greenland (stednavne udvalget for grønland) between 1934 and 1983 was formerly held by the danish geodetic institute, now part of kort & matrikel styrelsen (kms). this documentation included the work of the stednavneudvalget sub-committees and an almost complete set of the minutes of the meetings of the place name committee. on 1 jan uary 1984 the responsibility for place names in greenland was trans ferred to oqaasiliortut / grønlands sprognævn in nuuk, green land, which includes nunat aqqinik aalajangiisartut / grønlands stednavnenævn; the latter institute now holds the archives of the former place name committee for greenland. rgs report archive: royal geographical society (rgs), 1 kensington gore, london sw7 2ar, uk. the royal geographical society is the most important geographical organisation in england, and acts as an adviser to expeditions planning journeys to all parts of the world. it holds a large collection of unpublished expedition reports from 1965 onwards. asher, g.m. 1860: henry hudson the navigator. hakluyt society 1. series 27: ccxviii + 292 pp. backlund, h.d. & malmquist, d. 1932: zur geologie und petro graphie der nordostgrönländischen basaltformation. i: die basische reihe. meddelelser om grønland 87(5), 61 pp. backlund, h.g. 1930: contributions to the geology of northeast greenland. meddelelser om grønland 74(11), 207–296. backlund, h.g. 1932: das alter des ‘metamorphen komplexes’ von franz josef fjord in ost-grönland. meddelelser om grønland 87 (4), 119 pp. balchen, b. 1958: come north with me, 318 pp. new york: e.p. dut ton. bang, o. 1944: blant fangstfolk og bikkjer i eirik raudes land, 128 pp. oslo: kamban forlag. banks, m. 1955: commando climber, 240 pp. london: j.m. dent & sons. banks, m. 1957: high arctic, the story of the british north green land expedition, 276 pp. london: j.m. dent & sons. barr, w. 2010: the arctic voyages of louis-philippe-robert, duc d’orléans. polar record 46, 21–43. bartholomew, j.g. 1920: the times survey atlas of the world, 112 pla tes. london: the times. bartlett, [r.a.] ‘bob’ 1934: sails over ice, 301 pp. new york: charles scribner’s sons. bartlett, r.a. & bird, j. 1931: the bartlett east greenland expedition. geographical review 21, 398–414. battle, w.r.b. 1952: contributions to the glaciology of north east greenland 1948–49 in tyrolerdal and on clavering ø. med del elser om grønland 136(2), 28 pp. bay, e. 1896: geologi. meddelelser om grønland 19(6), 149–187. bearth, p. 1954: eindrücke aus den werner-bergen. berge der welt 9, 167–168. bearth, p. 1959: on the alkali massif of the werner bjerge in east green land. meddelelser om grønland 153(4), 63 pp. bendix-almgreen, s.e., clack, j.a. & olsen, h. 1988: upper devo nian and upper permian vertebrates collected in 1987 around kejser franz joseph fjord, central east greenland. rapport grøn lands geologiske undersøgelse 140, 95–102. bengtsson, m. 1927: ene med dyr og mennesker. et aar i scoresby sund, 229 pp. københavn: steen hasselbalchs forlag. bennet, d.j. 1959: climbing in the staunings alps, greenland. scot tish mountaineering club journal 26, 311–324. bennet, d.[j.] 1969: the staunings alps again. scottish mountain eering club journal 29, 146–150. bennet, d.j. 1972: stauning alps – greenland, 120 pp. reading: ga ston’s alpine books, west col productions. bennike, o. 1998: pingoer. grønland 1998, 62–68. bennike, o. & weidick, a. 1999: observations on the quaternary geology around niooghalvfjerdsfjorden, eastern north green land. geology of greenland survey bulletin 183, 57–60. bennike, o., mikkelsen, n., forsberg, r. & hedenäs, l. 2006: tup piap qeqertaa (tobias island): a newly discovered island off northeast greenland. polar record 42, 309–314. bennike, o., sørensen, m., fredskild, b., jacobsen, b.h., böcher, j., amsinck, s.l., jeppesen, e., andreasen, c., christiansen, h.h. & humlum, o. 2008: late quaternary environmental and cultural changes in the wollaston forland region, northeast greenland. advances in ecological research 40, 45–79. bennike, o., mikkelsen, n. & forsberg, r. 2009: tobias ø. grønland 2009, 18–25. berthelsen, c., mortensen, i.h. & mortensen, e. (eds) 1989: kalaallit nunaat grønland atlas, 130 + 24 pp. pilersuiffik. bierther, w. 1941: vorläufige mitteilung über die geologie des östlichen scoresbylandes in nordostgrönland. meddelelser om grøn land 114(6), 20 pp. bird, c.g. & bird, e.g. 1941: the birds of north-east greenland. ibis 14(5), 118–161. bistrup, h. 1924: ‘teddy’s sidste togt, 68 pp. københavn: g.e.c. gads forlag. bjerre, j. 1980: sirius. danmarks slædepatrulje i nordøstgrønland, 196 pp. københavn: forlaget komma. björck, s., wohlfarth, b., bennike, o., hjort, c. & persson, t. 1994: revision of the holocene lake sediment based chronology and event stratigraphy on hochstetter forland, ne greenland. bo reas 23, 513–523. bjørnbo, a.a. 1911: cartographia grønlandica. meddelelser om grøn land 48, 332 pp. blom, i. 1973: kampen om eirik raudes land. pressgruppepolitikk i grønlandsspørsmålet 1921–31, 438 pp. oslo: gyldendal. blosseville, j. de 1834: carte d’une parti de la côte du groënland oriental, reconnue par le brig la lilloise, sous le commandement der mr. j. de blosseville en 1833. bulletin de la société de géo graphie iie serie, no. 12, 400 only. bobé, l. 1915: lourens feykes haans kursskrifter for besejlingen af grønland, særlig diskobugten (1719). det grønlandske selskabs aarskrift 1915, 41–49. bobé, l. 1921: grønland i 200-aaret for hans egedes landing. be skrivelse af distrikterne i sydgrønland, julianehaab distrikt. distrikts historie. meddelelser om grønland 61, 537–577. bobé, l. 1928: early exploration of greenland. in: vahl, m. et al. 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95–102. brun, e. 1966: østgrønlands-overenskomsten. tidsskriftet grønland 1966, 127–136. bryan, k. 1959: the crossing to mestersvig. scottish mountaineering club journal 26, 324–329. buess, h.p. 1953: fahrten in der nunatakzone. berge der welt 8, 205–217. burch, s. 2003: gronau nunatakker range, first ascents. american alpine journal 2003, 282–283. bütler, h. 1954: die stratigraphische gliederung der mitteldevonischen serien im gebiete von kap franklin am kejser franz joseph fjord in zentral-ostgrönland. meddelelser om grønland 116(7), 126 pp. cabot, d., goodwillie, r. & viney, m. 1988: irish expedition to north-east greenland 1987, 150 pp. dublin: barnacle books. callomon, j.h. 1970: geological map of the carlsberg fjord – fossil bjerget area. meddelelser om grønland 168(4), 10 pp. callomon, j.h. & birkelund, t. 1980: the jurassic transgression and the mid-late jurassic succession in milne land, central east green land. geological magazine 117, 221–226. campbell, k.c. 1964: trinity 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greenland 1962. grønland 1963, 321–332. wright, g.t. 1957: cambridge university and marlborough college expedition to north-east greenland, 1956. polar record 1957, 349–350. zweifel, h. 1958: geologie und petrographie von nathorsts land (ne-grönland). meddelelser om grønland 160(3), 94 pp. 363 364 365 lynn ø anna bistrup fjelde leffingwell nunatak s k a l l i n g e n gnejsnæs bourbon ø pariserøerne nordmarken stormlandet ambolten brønlunds grav lægervallen gamle jim øer bloch nunatakker kap georg cohn iskap nakkehoved finderup land prinsesse dagmar ø prinsesse thyra ø 24˚ 20˚ 16˚ 12˚ 24˚28˚ 20˚ 16˚ 82˚ 81˚ 80˚ 79˚ 78˚ 81˚ 80˚ 79˚ 78˚ prinsesse ingeborg halvø kap jungersen k r o n p r i n s c h r i s t i a n l a n d m at h ild e a lp er pr in se ss e e lis ab et h a lpe r p ri n se ss e c ar ol in e h er tu ge n af o rl éa ns la nd gamma ø n ord ve stk lint a m d r u p h ø j l a n d v a n d r e d a l e ns j æ l l a n d f j e l d e n o r s e m a n d a l østkap nørre biland nørreland stigbøjlen h o l m l a n d hovgaard ø g r e n e n k i l e n a m d r u p l a n d l a m b e r t l a n d prinsesse margrethe ø f l a d e i s b l i n k nordostrundingen kap koefoed kap philippe kap montpensier depotnæsset station nord m y l i u s e r i c h s e n l a n d sophus müller næs 50 km henrik kröyer holme eskimonæs kap adolf jensen kap anna bistrup kap h.n. andersen north greenland erik s. henius land schnauder ø sydkap kap nansen k r o n p r i n s f r e d e r i k l a n d île de france qeqertaq prins henrik • j. c. c h r i s t e n s e n l a n d nørre mellemland søndre mellemland garde nunatakker hammeren tuppiat qeqertaa tobias ø east greenland bildsøe nunatakker alabama nunatak troldehaven danske øer franske øer norske øer kap amundsen moltke nunatakker kap amélie d a n m a r k f j o r d sk jo ld un ge el v antarc tic bugt hvidebugt hekla sund dijm phn a s und r o m er sø h a g e n f j o r d h a g e n b r æ i n d e p e n d e n c e f j o r d fyn sø centrumsø g r æ s e l v n i o g h a l v f j e r d s f j o r d e n z a c h a r i a e i s s t r ø m m armorvigen h anseeraq fjord sp al te gl et sc he r blåsø i n g o l f f j o r d blæsebræ wandel hav jøkelbugten map 1 3 5 4 2 1 366 lille koldewey stormlandet h er tu ge n a f o rl éa ns la nd søndermarken okselandet ymer nunatak d r o n n i n g l o u i s e l a n d kap steensby kap marie valdemar o rléans sundnordmarken depotnæsset eventyrfjelde k o n g w ilh el m l an d haystack rechnitzer land søndre mellemland daniel bruun land store koldewey gamma ø zebra klippe carlsbergfondet land trums ø adolf s. jensen land hochstetter forland shannon kap rink kap pansch pendulum øer kap david gray kuhn ø hvalrosø a. p. olsen land hold with hope home forland wollaston forland r ev et steno land payer land c. h. ostenfeld land kap berlin kap bremen knudshoved ole rømer land strindberg land bartholin land grejsdalen a.b. d rac hmann gletsch er budolfi isstrø m kap broer ruys jordanhill nanok ø teufelkap kap alf trolle daneborg j. l. mowinckel land brit annia gletsch er admiralty gletsch er b o rgjøkel borgfjorden sælsøen fl ad eb ug t penthievre fjo rd d o v e b u g t bessel fjord l . b i s t r u p b r æ bredefjord a rdencaple fjord g ra nd jea n fjord fligely fjo rd lindeman fjord claveringstrædet moskusoksefjord w ordie g letscherw a l t e r s h a u s e n g l e t s c h e r g eo lo gfjord d r o n n i n g m a r g r e t h e ii l a n d g e r m a n i a l a n d h u d s o n l a n d a n d r é e l a n dlouise boyd land bass rock lille pendulum alabama sabine ø kap börgen th. thomsen land s t o r e b æ l t so ra ne rb ræ en hochstetterbugten gael hamke bugt jackson ø kap stosch mackenzie bugt hamberg gletscher ado lf hoel gletscher krumme langsø l o c h fyn e eskimonæs finsch øer young sund c l a v e r i n g ø kap amélie kap philippe kap montpensier troldehaven bildsøe nunatakker s k æ r f j o r d e n k of oe d h an se n br æ danmark havn danmarkshavn 32˚ 77˚ 76˚ 75˚ 74˚ 28˚ 24˚ 20˚ 28˚ 20˚ 16˚ 77˚ 76˚ 75˚ 74˚ s t o r s t r ø m m e n alabama nunatak pendulumstrædet n o rd fjo rd g erard de geer gletscher arnold escher land île de france qeqertaq prins henrik • kap copeland ad astra iskappe 100 km map 2 3 5 4 2 1 367 grejsdalen gunnar andersson land y m e r ø s u e s s l a n d t r a i l l ø geographical society ø bontekoe ø kap mackenzie kap mcclintock kap wijkander kap bennet broch øer kap parry kap graah gletscherland kap simpson kap young mestersvig s t a u n i n g a l p e r olympen gurreholm sydkap bjørneøer h ä s i b j e r g e i s m a r k e n frænkel land r e n l a n ded va rd b ay d al storø m i l n e l a n d wat k i n s b j e r g e g å s e l a n d danmark ø gåsepynt l i v e r p o o l l a n d murray ø kap greg kap tobin savoia h alv ø v o l q u a r t b o o n k y s t k l i t d a l kap hewitt kap jones kap hodgson rathbone ø kap brewster steward ø manby halvø kap ryder hold with hope kap broer ruys andrée landlouise boyd land j.l. mowinckel land g a u s s h a l v ø ella ø l y e l l l a n d s c o r e s b y l a n d goodenough land petermann bjerg cecilia nunatak n a t h o r s t l a n d h i n k s l a n d c h a r c o t l a n dalfabet nunatakker brune nunatakker royston nunatakker graben land canning land j a m e s o n l a n d turner ø kap hope flakkerhuk kap leslie kap ewart kap dalton henry land c hr is ti an iv g le ts ch er kap franklin shackleton bjerg rigny bjerg hall bredning j.p. koch fjeld constable pynt illoqqortoormiut • scoresbysund gronau nunatakker wager nunatakker paul stern land b l o s s e v i l l e k y s t kap stewart f r a n z j o s e p h f j o r d geologfjord n o rd fjo rd k e j s e r kempe fjo rd isfjord n o rd en sk iöld gl hamberg gletscher concordia plads a lp ef jo rd furesø vega sund mountnorris fjord cambridge bugt fl em in g fjo rd n at ho rs t fjord c ar ls be rg fjord lillefjord d eichmann fjord v ik in gebugt b redegletscher charcot havn g e i k i e p l a t e a u knækket m agga d an g letscher g å s e f j o r d fønfjord sn es un d r ød ef jo rd s c o r e s b y s u n d barclay bugt ve stf jord gletsc he r gåseg let sch er vest fjord rolige bræ r ypefjo rd v in du e g letscher flyve rfjord ø f j o r d s c h u c h e r t f l o d mackenzie bugt f o s t e r b u g tdusén fjord sofia sund vega sund torvgletscher d endritgletscher k nighton fjord k o n g o s c a r f j o r d davy sund moskusoksefjord h u rry in le t st eno bræ narhvalsund harefjord syd b ræ daugaard jensen glet sc he r n o r d v e s t f j o r d 3 5 4 2 1 map 3 32˚ 72˚ 71˚ 70˚ 28˚ 24˚ 28˚ 24˚ 73˚ 72˚ 71˚ 70˚ 75 km in front pocket map 4: place name map of northern east greenland, 1: 1 000 000. a.k. higgins (2010). in back pocket map 5: place name map of stauning alper, 1: 150 000. a.k. higgins (2010). 368 geological survey of denmark and greenland bulletin 4, s 2003, pp 13-16 13 gestco is an acronym for european potential for the geological storage of co2 from fossil fuel combustion. the project formed part of the energie programme of the european union 5th framework and was concluded in 2003. the geological survey of denmark and greenland (geus) led the project, with the national geological surveys of belgium, france, germany, greece, the netherlands, norway and uk as research partners (fig. 1). the primary goal of the gestco project was to determine whether the geological storage of carbon dioxide captured at large industrial plants is a viable method of reducing greenhouse gas emissions in europe (christensen 2000; gale et al. 2001; christensen & holloway 2003). this was evaluated by a series of case studies that assessed the co2 storage potential of saline aquifers, geothermal reservoirs, coal seams and oil and gas reservoirs. the case study approach was used so that currently available, largely theoretical, generic information could be applied to real geological situations. in addition aspects of safety and environment, conflicts of using underground space and public and stakeholder perception were evaluated. secondary goals of the gestco project were to establish an inventory of major co2 point sources in europe and a decision support system (dss) to serve as an economic analysis tool for co2 storage in europe. inventory of large co2 point sources major industrial sources of co2 in the participating countries were identified and compiled into a database. in almost all countries, the major sources of co2 are power plants, fig. 1. map of europe showing the participating countries and offshore study areas included in the gestco project. geological survey of denmark and greenland bulletin 4, 13–16 (2004) © geus, 2004 assessing the european potential for geological storage of co2: the gestco project niels peter christensen and michael larsen 14 integrated steel plants, refineries/petrochemical complexes and cement works. the exception is norway, where many of the major sources of co2 are generators at offshore oil and gas fields. the location and details of the sources of co2 are compiled into a geographic information system (gis) enabling qualified search routines. in denmark, the annual emission of greenhouse gases is close to 60 mt of which approximately half originates from fossil fuel combustion related to power and heat generation. major co2 point sources were identified based on yearly reports to the danish energy authority. these point sources alone contribute 29 mt co2 of the total co2 emission in denmark. the largest single source is the coal-fired power plant asnæsværket in kalundborg, with an average yearly emission of 5.8 mt co2 in the period 1994–1999. considering co2 sequestration from asnæsværket could thus account for approximately half of the greenhouse gas reductions required for denmark in the kyoto agreement (larsen et al. 2003a). european storage capacities underground storage capacities in the case study areas were evaluated by seismic mapping, analysis of well logs and reservoir simulation. the results are summarised in christensen & holloway (2003). the major part of the mapped storage capacity was related to deep saline aquifers in onshore and nearshore sedimentary basins in denmark, germany, southern uk and northern france. in the netherlands and belgium the storage potential is primarily related to exhausted gas fields and coalmines. a huge potential exists in aquifers offshore norway, and it is likely that very large additional offshore aquifer potential exists in british and danish sectors of the north sea. the greek storage potential is composed of aquifers as well as a few hydrocarbon fields. significant storage capacity is related to the gas and oil fields of northern europe, particularly in the north sea and onshore in the netherlands and germany. geological storage capacities in denmark the potential storage capacity in denmark was evaluated through case studies of onshore and nearshore saline aquifers, and hydrocarbon fields of the danish north sea sector. deep saline aquifers large sedimentary basins of late palaeozoic – cenozoic age are present in denmark and provide a potential for co2 storage. in the onshore or nearshore danish area the reservoir units comprise porous sandstone layers of the lower triassic bunter sandstone formation / skagerrak formation, the upper triassic – lower jurassic gassum formation, the middle jurassic haldager sand formation, and the upper jurassic lower cretaceous frederikshavn formation. mapping and initial description of these units has been undertaken in the search for hydrocarbons and geothermal reservoirs (cf. nielsen et al. 2004, this volume). the gestco aquifer study was focused on sandstone formations within a depth range of 900–2500 m, i.e. between the depth required for co2 to become a dense fluid and the depth below which reservoir quality typically deteriorates due to diagenetically induced reduction of porosity and permeability. the total storage capacity of unconfined aquifers in denmark has been estimated to be 47 gt of co2, although only a small part of the volume was related to structural closures (holloway et al. 1996). in order to gain public and 15 political acceptance, structural traps are considered essential, at least initially, when considering storage onshore denmark, and consequently the gestco study was focused on eleven large structures (fig. 2; table 1). these structures were mapped from seismic surveys and evaluated using data from existing deep wells to assess the storage potential (larsen et al. 2003b). based on many years of experience from aquifer storage of natural gas in denmark, germany and france, it is assumed that 40% of the total pore volume within a trap could be filled with co2. this effective storage capacity will depend on a number of parameters including the geometry of the trap (e.g. difference in height between top point and spill point), the number of injection wells, injection rates and reservoir characteristics. the initial calculations carried out in the present study suggest that the eleven structures alone may provide storage for at least 16 gt of co2 (table 1; larsen et al. 2003b). note that almost two thirds of the calculated aquifer storage capacity is present in the thisted/legind structure. as well as a proper reservoir, a tight cap rock is needed when considering underground storage of co2. geological formations with good sealing properties are lacustrine and marine mudrocks, evaporites and carbonates. in denmark the most important sealing rock type is marine mudstone, which often forms units several hundred metres thick and is present at several stratigraphic levels. in addition to the primary cap rock, chalk of late cretaceous – danian age forms a possible secondary seal in most of the danish area. the sealing effect of the chalk is dependent on chemical reactions between dissolved co2 and the carbonate rock. detailed site surveys will be needed in order to test the integrity of the seal at any future storage site. storage in oil and gas fields although the potential storage capacity of deep saline aquifers is many times greater than that of hydrocarbon structures, there are some distinct advantages of using depleted hydrocarbon fields as storage sites. first, the hydrocarbon fields have proved their capability to retain fluids and gases, in many cases for millions of years. secondly, the reservoir is well understood due to intensive data gathering before and during the productive life of the field, and finally infrastructure for the production and transport of fluids and gases is already in place. with some modifications, this infrastructure may often be re-usable for delivery and injection of co2 for storage. the reserve figures for 14 chalk fields and three sandstone fields of the danish sector were included in the gestco project (christensen & holloway 2003). these comprise detailed estimates of the expected ultimate (initial) reserves, and rounded figures for low and high case reserves as given by the danish energy authority (dea 2002). the storage capacity of hydrocarbon reservoirs is calculated from the underground volume of ultimately recoverable oil or gas. the calculation assumes that the entire underground volume of recoverable hydrocarbons can be replaced by co2. for gas reservoirs, this is a straightforward assumption, since most gas reservoirs are of a closed nature. formation water (the aquifer) does not significantly replace the drained gas during fig. 2. map showing the position and outline of the eleven structural closures mapped in the danish aquifer case study of the gestco project. black dots indicate the position of deep exploration wells used in the evaluation of the reservoirs. modified from larsen et al. (2003a). 16 the producing field life. for oil reservoirs, it is assumed that the amount of co2 that can be stored in the reservoir is approximately 30% of the oil initially in place. since the ultimate recovery (ur) of most oil fields (as initially reported) also approximates 30–35%, the storage capacity can be approximated by the initial proven reserves (or ur). in addition to the volumetric estimates special concern is needed when considering storage of co2 in chalk, as chemical and physical reactions are likely to occur between co2 in solution and carbonate rocks. based on the above assumptions, the danish storage capacity for existing oil and gas fields is estimated to be 629 mt co2. of this, 452 mt can replace natural gas, while 176 mt can replace oil (christensen & holloway 2003). however, all of the investigated fields of the danish north sea are in the production phase, and co2 injection will probably not be possible in the near future unless applied through enhanced oil recovery (eor) operations. the eor option of the north sea is currently under investigation and preliminary results have been presented by markussen et al. (2003). the vision of the project is to capture co2 from the danish power plants and export it through an extensive pipeline system to the offshore industry. implementation of the eor technology would have great impact on the lifecycle of danish oil and gas production. conclusions the inventory of major point sources of co2 and the geological storage potential mapped in the gestco project indicate that the eight european countries could make a significant impact on their national co2 emissions by capturing the emissions from a relatively small number of the largest point sources and storing them underground. in denmark, mapping of geological structures suitable for underground storage of co2 suggests that enough storage volume is present within the subsurface to store several hundred years of total co2 emissions from danish industry and power production. detailed site surveys and risk analysis, including long term monitoring, are needed to validate these storage capacities and should be the focus of future studies. acknowledgements the project formed part of the energie programme of the european union 5th framework programme for research & development, project no. enk6-ct-1999-00010, and was 50% funded by the programme. in addition to the principal research partners, valuable contributions to the project were made by the flemish institute for technological research (vito, belgium), public power corporation of greece, compagnie française de geothermie (cfg), danish oil and natural gas company (dong), ce-transform (the netherlands) and the tyndall centre (uk). further contributions to the project were given by beb (now empg), bp, danish energy authority, gaz de france, iea greenhouse gas r&d programme, norsk hydro, norwegian petroleum directorate, shell, statoil, total-finaelf, the uk department of trade and industry, and vattenfall. references christensen, n.p. 2000: the gestco project: assessing european potential for geological storage of co2 from fossil fuel combustion. in: williams, d. et al. (eds): proceedings of the fifth international conference on greenhouse control technologies (ghgt-5), 261–265. cairns, australia: csiro publishing. christensen, n.p. & holloway, s. 2003: geological storage of co2 from combustion of fossil fuel. european union fifth framework programme for research and development, summary report, 34 pp. copenhagen, denmark: geological survey of denmark and greenland. dea 2002: oil and gas production in denmark 2001, 98 pp. copenhagen, denmark: danish energy authority. gale, j., christensen, n.p., cutler, a. & torp, t.a. 2001: demonstrating the potential for geological storage of co2: the sleipner and gestco projects. environmental geosciences 8, 160–165. holloway, s. et al. 1996: the underground disposal of carbon dioxide. joule ii project no. ct92-0031, summary report, 24 pp. keyworth, nottingham, uk: british geological survey. larsen, m., bidstrup, t. & christensen, n.p. 2003a: saline aquifer storage of co2 from major point sources – a danish case study. in: gale, j. & kaya, y. (eds): greenhouse gas control technologies. proceedings of the sixth international conference (ghgt-6), kyoto, october 2002, 657–661. larsen, m., bidstrup, t. & dalhoff, f. 2003b: co2 storage potential of selected saline aquifers in denmark. danmarks og grønlands geologiske undersøgelse rapport 2003/39, 83 pp. markussen, p., austell, j.m. & hustad, c.-w. 2003: a co2-infrastructure for eor in the north sea (cens): macroeconomic implications for host countries. in: gale, j. & kaya, y. (eds): greenhouse gas control technologies.proceedings of the sixth international conference (ghgt-6), kyoto, october 2002, 1077–1082. nielsen, l.h., mathiesen, a. & bidstrup, b. 2004: geothermal energy in denmark. geological survey of denmark and greenland bulletin 4, 17–26 (this volume). authors’ address geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. e-mail: npc@geus.dk << /ascii85encodepages false /allowtransparency false /autopositionepsfiles true /autorotatepages /all /binding /left /calgrayprofile (dot gain 20%) /calrgbprofile (srgb iec61966-2.1) /calcmykprofile (u.s. web coated \050swop\051 v2) /srgbprofile (srgb iec61966-2.1) /cannotembedfontpolicy /warning /compatibilitylevel 1.4 /compressobjects /tags /compresspages true /convertimagestoindexed true /passthroughjpegimages true /createjdffile false /createjobticket false /defaultrenderingintent /default /detectblends true /colorconversionstrategy /leavecolorunchanged /dothumbnails 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<feff4f7f752890194e9b8a2d5b9a5efa7acb76840020005000440046002065874ef65305542b8f039ad876845f7150cf89e367905ea6ff0c4fbf65bc63d066075217537054c18cea3002005000440046002065874ef653ef4ee54f7f75280020004100630072006f0062006100740020548c002000520065006100640065007200200035002e0030002053ca66f465b07248672c4f86958b555f3002> >> >> setdistillerparams << /hwresolution [2400 2400] /pagesize [595.000 842.000] >> setpagedevice geological survey of denmark and greenland bulletin 42, 2018, 39-64 39 sedimentology, geochemistry and reservoir properties of upper jurassic deep marine sediments (hareelv formation) in the blokelv-1 borehole, jameson land basin, east greenland morten bjerager, claus kjøller, mette olivarius, dan olsen and niels h. schovsbo the fully cored blokelv-1 borehole was drilled through upper jurassic strata in the central part of the jameson land basin, central east greenland. the borehole reached a total depth of 233.8 m with nearly 100% recovery of high-quality core. an extensive analytical programme was undertaken on the core; sedimentological interpretation and reservoir characterisation were based on facies analysis combined with conventional core analysis, bulk geochemistry and spectral gamma and density scanning of the core. the upper jurassic hareelv formation was deposited in relatively deep water in a slope-to-basin setting where background sedimentation was dominated by suspension settling of organic-rich mud in oxygen-depleted conditions. lowand high-density gravity-flow sandstone interbeds occur throughout the cored succession. about two-thirds of the high-density turbidite sandstones were remobilised and injected into the surrounding mud-rock. the resulting succession comprises nearly equal amounts of mudstones and sandstones in geometrically complex bodies. ankerite cementation occurs in 37% of the analysed sandstones in varying amounts from minor to pervasive. such ankerite-cemented sandstones can be identified by their bulk geochemistry where ca > 2 wt%, mg > 1 wt% and c > 1 wt%. the analysed mudstones are rich in al, fe, ti and p and poor in ca, mg, na and mn. the trace-metal content shows a general increase in the upper part of the core reflecting progressive oxygen depletion at the sea floor. the reservoir properties of the blokelv-1 sandstones were evaluated by both conventional core analysis and using log-derived porosity and permeability curves. the high-density turbidite beds and injectite bodies are a few centimetres to several metres thick and show large variations in porosity and permeability, in the range of 6–26 % for porosity and 0.05–400 md for permeability. individual sandstone units that are 1–7 m thick yield a net vertical reservoir thickness of 40 m with porosities of 15–26% and permeabilities of 1–200 md. heterolithic sandstone–mudstone units are generally characterised by poor reservoir quality with porosities of 2–14% and permeabilities of 0.1–0.6 md. keywords: hareelv formation, jameson land basin, upper jurassic, core description, facies analysis, reservoir quality, petrophysical logs ___________________________________________________________________________ geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. email: mbj@geus.dk © geus, 2018. geological survey of denmark and greenland bulletin 42, 39–64. available at: www.geus.dk/bulletin42 mailto:mbj@geus.dk http://www.geus.dk/bulletin42 4040 the upper jurassic succession in jameson land represents a marine seaway; the depositional environments recorded include oxfordian sandy shelf-edge and slope deposits in the north (olympen formation) and oxfordian–volgian base-of-slope and basinal mudstones and intercalated massive sandstones of the hareelv formation in the central and southern part (surlyk 2003). the upper part of the succession is represented by coarsefig. 1. geological map of jameson land, showing the location of the blokelv-1 borehole and selected outcrop localities. the map is based on the geus digital greenland geological map (1:500 000) and printed map series (1:100 000); only named rivers are shown. ; ; ; ; ; ; ca r l sb e r g f j o r d gø01_02_273_02_mbj_blokelv_sed gåseelv hareelv ugleelv blokelv ran unk ele lv hors ens fjo rd ra uk elv blokelv-1 jameson land liverpool land h urry in le t olympen fossilbjerget mikael bjerg katedralen fortet mbj art1 �g1 71°n 71°30'nmajor dyke/sill fault ice rivers (selected) quaternary hesteelv fm raukelv fm hareelv fm, salix dal mb olympen fm fossilbjerget fm/pelion fm neill klinter gp kap stewart gp triassic basement hareelv fm, sjællandselv mb hareelv fm, katedralen mb 70°30'n 23°w 24°w 22°w 20 km 71°'n 71°30'n ?devonian s c o r e s b y s u n d 41 grained shelf-edge clinothems of the raukelv formation in southern jameson land (surlyk & noe-nygaard 2005). the succession has attracted numerous outcrop studies particularly in the northern area (larsen & surlyk 2003; bruhn & surlyk 2004), but also in the east where the oxfordian–kimmeridgian is exposed (surlyk et al. 2007) and in the south where volgian strata crop out (surlyk & noe-nygaard 2005). the fully cored blokelv-1 borehole was drilled in upper jurassic strata in the central part of the jameson land basin (fig. 1), as part of geus’ activities in the project entitled ‘petroleum geological studies, services and data in east and northeast greenland’ (bojesen-koefoed et al. 2009; bjerager et al. 2018a, this volume). the core thus presents new data in an area that has previously been poorly known. the core spans the middle oxfordian to lowermost volgian interval (alsen & piasecki 2018, this volume). the katedralen member of the hareelv formation overwhelmingly dominates the section, comprising intercalated mudstones, sandstones and heteroliths from 233.8 (td, total depth) to 10.1 m. the lowermost part of the core (233.8–208.13 m) is of middle oxfordian age and thus correlates temporally with the uppermost part of the olympen formation. indeed, the bioturbated mudstones that occur locally in the lower 10 m of the core are reminiscent of the hades member of the olympen formation but due to the limited data they are referred to the hareelv formation. the sandstone-dominated sjællandselv member of the hareelv formation is recorded in the uppermost 10.1 m of the core, corresponding to the mapped unit outcropping at the borehole locality on the 1:100 000 geological map (fig. 1). the hareelv formation has been described as “…a giant sedimentary injection complex” (surlyk et al. 2007, p. 1), being characterised by slope gully-fill and basin-floor sandstone units that post-burial were remobilised and injected into the enveloping black basinal mudstones (surlyk & noe-nygaard 2001, 2003; surlyk et al. 2007). based on outcrop studies, all sandstone units in the katedralen member were considered to be post-depositionally remobilised, with obliteration of all primary sedimentological structures (surlyk et al. 2007). a consequence of the injected nature of the sandstone units is that interpretation of the primary depositional processes of the hareelv formation sandstones is challenging. in addition, the three-dimensional geometries of the sandstone bodies are complex and unpredictable. the blokelv-1 core is particularly valuable in contributing to the sedimentological and stratigraphic understanding of the hareelv formation, firstly since it permits an evaluation of a continuous and complete succession and secondly since primary sedimentary structures are commonly preserved. this facilitates the recognition of a wide spectrum of facies for which depositional processes can be evaluated. the core also reveals a wide range of post-depositional processes including water escape, slumping, fluidisation, injection and brecciation. some of these processes were essentially syndepositional, whilst others occurred after significant burial and compaction. this study presents the sedimentological description combined with an analysis of reservoir parameters (porosity and permeability), bulk inorganic geochemistry and petrophysical studies of core material from the upper jurassic of the blokelv-1 borehole in the jameson land basin. combination of these methods provides a valuable dataset that not only contributes to the understanding of the depositional environment and the postdepositional features of the hareelv formation but also facilitates an assessment of the reservoir quality of the sandy intervals. this paper is supported by, and complements, the remaining papers in this bulletin that focus on the stratigraphy and basin architecture (alsen & piasecki 2018; bjerager et al. 2018b), potential source rocks (bojesen-koefoed et al. 2018), diagenesis (olivarius et al. 2018a), provenance (olivarius et al. 2018b), uplift history (green & japsen 2018), and igneous intrusions (larsen 2018). methods the blokelv-1 core was logged sedimentologically at a scale of 1:20 (figs 2–4). overview core photographs were produced for the entire core, and selected core intervals were slabbed and imaged at high resolution. a total of 48 thin sections were made for microfacies characterisation and petrography, with respect to both the sedimentology presented here and the diagenetic study presented in olivarius et al. (2018a, this volume). all analyses were performed at geus. geochemistry. major and trace elements were analysed on 42 samples selected from the different facies in the core. analyses were made by inductively coupled plasma emission spectroscopy (icp-es) and inductively coupled plasma mass spectroscopy (icp-ms). total carbon (tc) and total sulfur (ts) were measured by leco analysers and loss on ignition (loi) was measured after heating to 1000°c. all measurements were made on rock powder. a clean half-core sample of 1–2 cm length was used for each measurement. 4242 stage m id dl e up pe r lo w er up pe r vo lgi an o xf or di an ki m m er id gia n lo w er depth (m) 10 depth (m) 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 210 220 230 10 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 210 220 230 233.8 be be be 2.5 0 0 0 0 5 10 1510 201 2 3100 2003.0 filtered density log grtotal (api units) k2o (%) u (ppm) 20 th (ppm) gø04_02m_453_mbj_blokelv_sed b a mudstone heterolith (mudstone/sandstone) sandstone, gravity flow sandstone, injectite lithology igneous intrusion be ripple cross-lamination wavy bedding slump sandstone intrusion small mudstone intraclasts bentonite large mudstone clast coalified wood belemnite ammonite bivalve analytical data point brachiopod structures, biota parallel lamination/bedding diffuse stratification clay si sand pebbl. mbj art1 �g 2 43 porosity, permeability, grain density. he-porosity and single point n2 gas permeability were measured on 35 samples. measurements were made on plugs drilled from the range of core lithologies: sandstone (25 horizontal plugs, two vertical (relative to bedding)), mudstone (five vertical), and igneous intrusions (three horizontal). the measurements follow the american petroleum institute (api) recommended practice for core analysis proce dure (api 1998). samples were drilled as 1 inch diameter cylindrical plugs using tap water as coolant. water and dissolved salt were completely removed by meoh cleaning in a soxhlet extractor. the samples were dried at 40°c for 72 hours before analysis. permeability was measured in a hassler core holder at a confining pressure of 400 psi (27.6 bar). the uncorrected gas permeability was measured by flowing nitrogen gas through a plug of known dimensions at differential pressures between 0 and 14.5 psi (0–1 bar). no back pressure was applied. the lower limit for permeability measurements is 0.05 md. the porosity was determined by subtraction of the measured grain volume from the measured bulk volume. the helium technique, employing boyle’s law, was used for grain volume determination, applying a doublechambered helium porosimeter with digital readout. the sample bulk volume was measured by submersion of the plug in a mercury bath using archimedes principle. grain density was calculated from the grain volume measurement and the weight of the cleaned and dried sample. precision of analytical data with reproducibility at the 68% level of confidence (+1 standard deviation) is 0.003 g/cm3 for grain density and 0.1% (p.u.) for porosity. the precision of permeability measurements is 15% in the range 0.05–0.1 md and 6% > 0.1 md. core logging. a set-up that allows simultaneous spectral gamma-ray and density measurements to be recorded was used for petrophysical logging of the core; the scanning speed was 1 cm min-1 and data retrieval occurred once per minute. the spectral gamma-ray analysis was carried out using two 15 cm crystals of sodium iodide activated with thallium (nai (tl)) and the bulk density was determined using a caesium source. the core sections, fitted together, were scanned sequentially so that the continuous core passed slowly beneath the scanner detectors. the density data were filtered to remove the large amplitude variations caused by missing core, crushed core or from gaps between the termination of one core section and the beginning of the next. the principle in the filtering process is a comparison between each individual data point and a running average calculation of the measurements. the procedure was as follows: step 1. establish a running average measurement calculated in a five-point data window according to: n d ravg i ik k i ∑ + −== 2 2 where di is the density value at depth i, and n is the number of data measurements in the five-point data window. the number of real data points (i.e. the value of n) in the window may vary from 5 to 0, depending on the number of missing data points. step 2. filtering procedure, according to: ravgi – di < 0.1g/ml in step 2, the data (di) with a difference from the average value larger than 0.1 g/cm3 were removed from the dataset. steps 1 and 2 were repeated five times, giving five filtering runs. the efficiency of each run was evaluated by comparing the number of data removed in the run. only a very small number of data measurements were removed in the last two filtering runs. in three intervals of the core, comparison was made between core photographs and density logs of the unfiltered data, filtering run #3 and filtering run #5 in order to validate the filtering procedure. in all cases, the filtering procedure was found to have been applied successfully such that data points were only filtered out in crushed core intervals and at gaps between core pieces. after filtering, the missing density measurements were linearly interpolated from neighbouring measurements to present a continuous log. calibration of the spectral gamma and density scanner was made by running in-house standards with known uranium (u), potassium (k) and thorium (th) concentrations and known densities. after calibration, the core scanner measurements of u, k and th were compared with concentrations of these elements measured geochemically in core plugs, and the scanner measurements facing page: fig. 2. generalised lithological log of the blokelv-1 core (geus 511101) showing the chronostratigraphic subdivision and corederived petrophysical logs (filtered density (g/cm3, total (grtotal) and spectral (k2o, th and u) gamma logs). the legend is also applicable to fig. 12. the location of the detailed sedimentological logs in figs 2 and 3 are shown by a and b, respectively. 4444 rescaled to fit the amplitude of core measurements (fig. 2). total gamma activity is reported in counts per second (cps). the following empirical relationship has been established between api units and the elemental concentrations. the relationship is not certified and should only be used as a rough guideline: total gamma ray (api units) = 8 × u (ppm) + 4 × th (ppm) + 16 × k (%) for presentation purposes, a median filtering of 25 cm was run on the density data to smooth the variation. the blokelv-1 core was partially water saturated during the spectral gamma and bulk density scanning. seven core pieces were weighed before scanning, after scanning, and finally after drying in an oven at 60°c for at least 70 hours; the oven drying reduced the bulk density of the core pieces by 0.00–0.20 g/cm3. the findings indicate that the presented spectral gamma and bulk density results pertain to a water-saturated core. log-derived porosity and permeability. the bulk density values were converted to porosities in two steps: step 1. porosities were calculated according to: φ = (ρma – ρb)/(ρma – ρf ) where: φ = porosity, ρma = matrix (or grain) density, ρb = formation bulk density (log value), ρf = density of the fluid saturating the rock. this calculation assumed an average matrix density calculated from core analysis of 2.68 g/cm3 and 100% saturation of fluid with assumed density of 1.0 g/cm3. this porosity calculation results in negative porosities from intervals with bulk densities >2.68 g/cm3 and in much higher porosities in clay-rich intervals than measured by conventional core analysis. matrix densities may vary due to clay volume, degree and type of carbonate cementation and the occurrence of igneous rock, all factors that are relevant to the blokelv-1 core. a second step was thus applied in order to calibrate measured porosities (on plugs) with logderived porosities following the method outlined by esbensen et al. (2015). step 2. a partial least squares (pls) regression analysis is made on a common dataset of core measurements and log readings (porosity from step 1), formation bulk density, k, th, u and total gamma (api). this analysis provides the best fit for calibration of core and log data and utilises all recorded corescanning parameters in the final porosity estimation. in a similar manner, core-scanning data were calibrated using the measured core permeabilities to derive a permeability log. the core scanning and derived porosity and permeability logs are combined with the sedimentological log in order to characterise the lithological units, e.g. variation in density in clean sandstone intervals. the spectral gamma-log traces were used to identify depositional trends in the mudstone-dominated intervals (see bjerager et al. 2018b, this volume). sedimentology the blokelv-1 core comprises a succession of intercalated mudstones and sandstones; nine sedimentary facies are recognised and briefly described below. laminated mudstone makes up 52% of the total thickness of the cored section; the heterolithic sandstone–mudstone facies represents 11%, and the massive and chaotic sandstones represent 15% and 21%, respectively (figs 2–4). subordinate facies include muddy sandstone and biomottled mudstone, each representing about 1% of the core section, and muddy shell beds and bentonite layers (<1%). the petrology and diagenesis of these sedimentary facies are described by olivarius et al. (2018a, this volume). laminated mudstone this facies consists of very dark grey to black mudstone (figs 3, 4), laminated at a millimetre to centimetre scale, and with a relatively high content of organic material (average total organic content (toc) c. 7%) and pyrite (total sulfur (ts) 1–8 wt%) (bojesen-koefoed et al. 2018, this volume). lamination is commonly defined by thin siltstone or very fine-grained sandstone pinstripe laminae (fig. 5a). at some levels, these siltstone and sandstone form laminae and lenses up to several millimetres thick (fig. 5b) and where abundant represent a transition to the heterolithic facies (see below). pyritised ammonites and thin-shelled buchiid bivalves dominate the macrofauna (alsen & piasecki 2018, this volume). trace fossils are very rare. extraformational coalified wood fragments (0.3–0.5 cm across) and plant remains facing page: fig. 3. sedimentary log of the interval 225.84–222.07 m of the blokelv-1 core with accompanying photographs of slabbed core of this interval. the represented facies comprise (from below): heterolithic sandstone–mudstone, black laminated mudstone with local beds of biomottled mudstone, and massive sandstone. 45 10 cm toptop cl. si. vf f m c vc f m c sand pebbl. co. 223 depth (m) 224 225 py py py py py py 15° 15° basebase laminated mudstone: hemipelagites, low-density turbidites mudstone-clast breccia mudstone sandstone heterolith: low-density turbidites, hemipelagites massive sandstone: high-density turbidites facies, processes ripple cross-lamination wavy bedding slump, disturbed bedding small mudstone intraclasts sandstone intrusion large mudstone clast bioturbation parallel lamination/bedding diffuse stratification chaotic sandstone, mudstone-clast breccia: injectites pyrite dip of bedding/layering heterolithic sandstone/mudstone gø04_02m_462_mbj_blokelv mbj art1 �g 3 4646 10 cm toptop basebase 36 cl. si. vf f m c vc f m c sand pebbl. co. 37 38 39 py 20° 50° 25° 70° gø04_02m_461_mbj_blokelv mbj art1 �g 4 depth (m) fig. 4. sedimentary log of the interval 38.98– 35.31 m of the blokelv-1 core with accompanying photographs of slabbed core of this interval. the represented facies comprise (from below): massive sandstone, black laminated mudstone with sandstone injectites, slumped sandstone with sandstone injectites, injected chaotic sandstones and mudstone-clast breccias and laminated mudstone with few thin sandstone injectites. for legend, see fig. 3. 47 are present in places. although the organic content was mainly derived from marine algae, there is a significant terrigenous component in the lower part of the core that decreases upward (bojesen-koefoed et al. 2018, this volume). the facies occurs throughout the core in units that range from a few millimetres to more than 10 m in thickness (figs 2–4). this laminated mudstone facies is interpreted to have mainly settled out from suspension, accumulating on the sea floor under oxygen-depleted conditions. the sharpbased siltstone and very fine-grained sandstone laminae are interpreted to record deposition from low-density turbidity currents. biomottled mudstone the facies consists of grey to dark grey silty or sandy mudstone with a weakly biomottled fabric; chondrites and planolites traces occur locally. coalified wood and plant fragments up to a few millimetres in size are common. the facies is only recognised locally in the lower levels of the core (231.8–231.0; 224.8–223.8 m, fig. 3). the mudstone is interpreted to have been deposited from suspension and/or from distal low-density turbidity currents; dysoxic oxygen conditions prevailed at the sea floor permitting colonisation by a limited infauna. muddy sandstone this facies consists of dark grey muddy sandstone with abundant sand-sized mudstone intraclasts and minor larger mudstone clasts up to several centimetres across (fig. 5g). the facies is recognised in four discrete units, 0.2–0.4 m thick, at 163 m, 146.5 m, 49 m and 41 m in the core. some units have irregular bases and sharp tops whereas others have flat, subparallel boundaries and show internal undulating or planar stratification. units with irregular bases and tops are also present. this facies cannot be attributed to a single process. beds showing sub-parallel boundaries are interpreted as debris-flow deposits, where consolidated mudstone clasts may have been derived from slumping or were incorporated as rip-up clasts during flow. certain units, however, show highly irregular boundaries and are probably injected sediment, perhaps representing sand injected into differentially consolidated mud that was incorporated as both matrix and clasts. heterolithic sandstone–mudstone the facies consists of alternating beds and laminae of grey fine-grained sandstone and dark grey to black mudstone (fig. 5c, d); the facies commonly forms fining-upwards units up to c. 1 m thick. the laminae/beds may show a concentration of mudstone intraclasts at the base, grading upwards into medium-grained sandstone (fig. 6a). coalified wood fragments and plant remains are common. thicker beds, 10–30 cm thick, commonly show an upward succession of sedimentary structures from planar or wavy lamination, through cross-lamination to parallel lamination. the cross-lamination may pass upwards into climbing ripple lamination, convolute and disrupted bedding, slump structures, and fluidisation structures (figs 2, 5d). individual beds and laminae have sharp and commonly erosional bases and their tops can be erosional, sharp or gradational (fig. 6a, b); they may be nongraded, normally graded or, less commonly, show inverse grading. diminutive flute casts or load casts may occur at bed bases. thick heterolithic intervals occur at 231–225 m, 191–186 m and 12–23 m; minor units occur at 168 m, 165 m, 153 m, 146 m, 82 m and 78 m (fig. 2). the sharp planar or erosional bed bases, the presence of normal grading and the succession of sedimentary structures within discrete sandstone beds or laminae are features indicative of deposition from waning turbulent flows. the sandstone component is thus interpreted as having been deposited from low-density turbidity currents (tb–d of bouma 1962); mudstone interbeds represent hemipelagic suspension deposits and/or deposition from the muddy tail of turbidity currents. the heterolithic facies resembles the so-called ‘tiger striped’ beds described in the literature and similarly interpreted as low-density turbidites of bouma types tb-e (e.g. lowe 1982; stow et al. 1982; guy 1992; surlyk & noenygaard 2001). the common occurrence of associated plant remains and coalified wood fragments, suggestive of a fluvial sediment source, and the inverse grading observed locally may indicate a hyperpycnal, flood-induced origin for some of these density flows (e.g. soyinka & slatt 2008; bhattacharya & maceachern 2009). massive sandstone this facies consists of grey, fineto medium-grained and moderately to well-sorted sandstones composed mainly of angular to subrounded quartz grains (figs 5e, f; see also fig. 8d). mica is common and locally dominant at the top of beds. individual sandstone beds range from 10 4848 mbj art1 �g 5 2 cm a b c d e f g h i 49 cm to several metres in thickness and have sharp lower and upper boundaries, in places forming amalgamated units up to 7 m thick (figs 2, 3). bed bases are commonly erosional and, where not amalgamated, upper boundaries are sharp, planar and succeeded by mudstone. the sandstones are generally structureless, but planar and gently undulating stratification may occur. water-escape structures and consolidation laminae are common in certain intervals. intraformational mudstone clasts in the fine sand to granule fraction are common to abundant in certain beds. the sand-grade mudstone intraclasts are dominantly rounded and evenly dispersed. the coarser mudstone clasts are commonly elongate and may occur evenly distributed through the bed, they may form a thin layer at the base (fig. 6b) or they may show an increase in abundance towards the top of the bed (fig. 8a). rare larger mudstone clasts, up to several centimetres long, occur as isolated floating clasts. sand injection features (dykes) may protrude from the top or from the base of the beds (i.e. extending upwards or downwards), and display features transitional to the chaotic sandstone described below. the massive sandstones are interpreted as high-density turbidites, an interpretation compatible with that of surlyk et al. (2007) for comparable sandstones in the exposed hareelv formation. the transporting gravity flows may have been initiated as slumps and sandy debris flows but subsequently evolved into high-density turbidity currents/ hyperconcentrated gravity flows (mulder & alexander 2001; larsen & surlyk 2003; bruhn & surlyk 2004). chaotic sandstone this facies consists of grey, fineto medium-grained and moderately to well-sorted sandstone. it is petrologically comparable to the massive sandstone facies being dominated by angular to subrounded quartz grains (fig. 8b); mica is common and may be dominant at the top of some sandstone bodies. individual units are 0.02–7 m thick in the bloklev-1 core (figs 2, 4). the facies has also been recorded from outcrop sections east of blokelv-1 where it forms units up to 50 m thick at hareelv and katedralen (fig. 1; surlyk & noe-nygaard 2001). in the blokelv-1 core, both the upper and lower boundaries of these sandstone units commonly truncate the adjacent mudstones or sandstones (fig. 6c). the truncation surface may be highly irregular or appear planar; truncation dip angles are up to 60°. ‘ptygmatically’ folded subvertical centimetre-scale sandstone dykes commonly protrude upwards from the sandstone units (fig. 5h). also present are water-escape structures such as dish-and-pillar structures in sandstones, and soft-sediment deformation structures in mudstones. certain intervals show common to abundant intraformational mudstone clasts, in the fine sand to granule fractions, as also observed in the high-density turbidite sandstone facies (figs 6c, 8c). large irregular and angular mudstone intraclasts are scarce and randomly distributed overall, but are locally common to abundant, representing a transition into the mudstone-clast breccia facies (fig. 8e). the facies is characterised by evidence of post-depositional deformation and/or remobilisation, and is interpreted to have been derived from the massive sandstones. the facies is termed ‘chaotic’ due to the complex, commonly irregular geometry of boundaries combined with the characteristic post-burial remobilisation features. the small ‘ptygmatically’ folded sub-vertical dykes were injected during early burial as they were folded during further burial and compaction of the mudstone (surlyk et al. 2007). large-scale liquefaction and injection of sand were commonly associated with brecciation of the adjacent compacted mudstone and large angular mudstone clasts were incorporated randomly in the injected massive sand (surlyk & noe-nygaard 2001; hurst et al. 2003; jonk et al. 2005; surlyk et al. 2007). remobilisation of the sands thus occurred at both shallow and deeper levels of burial. in some cases, an entire unit appears to have been completely liquefied and homogenised; in other cases, liquefaction was localised such that portions of the sand body retain primary depositional structures (fig. 5f). differentiation between this facies and the massive sandstones can be somewhat arbitrary based facing page: fig. 5. sedimentary facies (scale in i is applicable to all photographs). a: black laminated organic-rich mudstone (68.4 m); note the well-developed lamination and absence of bioturbation (also in b). b: dark grey organic-rich mudstone interbedded with abundant siltstone and fine-grained sandstone laminae (109.5 m). c: parallel-laminated heterolithic sandstone–mudstone (146.5 m). d: heterolithic sandstone–mudstone showing ripple cross-lamination, convolute bedding and fluidisation structures (224.8 m). e: massive sandstone; note the homogeneous, structureless fabric (223.5 m). f: sandstone (96.8 m) showing diffuse, undulating stratification, probably a combination of depositional (upper part) and injection structures (lower part). g: muddy sandstone with abundant sandsized mud clasts (146.7 m). h: sandstone dyke injected into black mudstone (232.7 m). i: heterogeneous fabric composed of interleaved mudstone-clast breccia and massive sandstone (71.2 m). 5050 on core study, being reliant on the limited view of bed boundaries in core; it is possible that some units referred to the massive sandstone facies are also injectites. mudstone-clast breccia this facies consists of angular mudstone clasts, millimetres to decimetres in size, in a fineto medium-grained sandy matrix (figs 4, 5i); the facies term is adopted from duranti & hurst (2004). both clastand matrix-supported fabrics are recorded; sorting of the clast fraction is variable but typically poor. the clasts may be concen2 cm mbj art1 �g 6 a b c d fig. 6. facies successions (scale in a is applicable to all photographs). a: sandstone-dominated interval of the sandstone–mudstone heterolithic facies showing amalgamated sharp-based (arrows) parallel-laminated sandstone turbidites (tb of bouma 1962); note the concentration of mudstone intraclasts in the basal part of the beds (189.7 m). b: interval (82.7 m) showing thin-bedded, sand-rich heteroliths (base and uppermost) interbedded with an inferred bentonite bed (7.5 cm thick) that is erosionally overlain by thin massive or graded sandstone beds rich in mudstone intraclasts. c: sandstone injectite containing abundant sand-grade mudstone debris, intruding upwards into parallel-laminated sandstone (152.8 m). d: sandstone injectite with a high-angle, sharp and irregular (cuspate) boundary to the host mudstone and containing elongate, highly angular rip-out mudstone clasts (84.5 m). 51 trated in irregular zones but where forming diffuse layers the mudstone-clast breccias commonly define steep dips (up to about 60°; fig. 4). the facies is well developed at 103 m, 71 m and 37 m. this mudstone-clast breccia facies has been recorded from outcrops of the hareelv formation in association with injected sands (surlyk & noe-nygaard 2001), yet it also resembles a subfacies of the lenticular massive sandstone facies in the underlying olympen formation, of inferred depositional origin (larsen & surlyk 2003). indeed, such mudstone-clast breccias are commonly reported in association with deep-water sandstones (e.g. johansson & stow 1995), and thus are not an unambiguous indicator of intrusion processes. in the blokelv-1 core, the facies occurs within units interpreted to have undergone post-depositional remobilisation and injection during burial, as indicated by the steeply dipping layering and the irregular distribution of mudstone clasts (fig. 4). the mudstone-clast breccias thus probably formed as the direct result of injection: consolidated mudstone clasts were detached from the host mudstone during sand intrusion and injection, and were incorporated into the invading sandy matrix (fig. 6d). it is possible, however, that some of the mudstone-clast breccias in the blokelv-1 core had a depositional origin as sediment gravity flows, analogous to those of the olympen formation (larsen & surlyk 2003), and were subsequently remobilised and involved in sediment intrusion. bioclastic mudstone and sandstone this facies is represented by a 15 cm thick unit at 83.5 m depth. it occurs within an interval (93–67 m) showing background high gamma-ray readings in mudstones, corresponding to a high content of uranium (>15 ppm; fig. 2); moreover, a discrete gamma-ray peak, indicating 2 cm gø06_658_mbj.eps mbj art1 �g7 fig. 7. core photograph of parallel-laminated mudstone abruptly overlain (arrow) by the heterogeneous shell-rich facies recognised only at 83.5 m; note the coalified wood fragment immediately above the base. the shelly sediment is composed of two discrete portions. the lower bioclastic bed (7 cm thick) shows sharp boundaries and is capped by a structureless mudstone lamina (0.5 cm thick). the upper portion, sharply overlying the mudstone (dashed line), comprises a stack of thin, graded sand-grade bioclastic layers (see text for detailed description and interpretation). 5252 gø06_621_mbj mbj fig8 0.5 cm a b c d e f 53 uranium levels over 50 ppm was registered within the facies unit. this distinctive shell-rich unit, recognised only at this level in the core, is composed of two discrete portions. at the base is a 7 cm thick shell bed with sharp boundaries (figs 7, 8f). it consists of abundant thinshelled bivalve fragments up to 1 cm long in a muddy matrix containing dispersed sand-grade quartz grains; a coalified wood clast (3 mm thick) occurs near the base. the shell fragments are oriented parallel to bedding. this shell bed has a diffuse internal wavy stratification defined by variation in the relative proportions of shells and matrix (fig. 8f). a structureless mudstone lamina, 0.5 cm thick, caps the muddy shell bed (fig. 7). the succeeding unit (c. 8 cm thick), erosionally overlying the mudstone cap of the lower muddy shell bed, comprises a stack of thin (1–3 cm), normally graded sand-grade bioclastic beds; the grains are dominated by bioclasts (bivalve fragments) with subordinate sandgrade mud clasts. the mud-rich nature of the basal shell bed, angular shell fragments often being supported by the mud matrix, is indicative of deposition from a debris flow; the thin mudstone cap may reflect loss of matrix cohesion at the upper boundary of the plastic flow. in contrast, the succeeding stack of graded, sand-grade bioclastic beds suggest deposition from a waning surging turbulent current. although the two portions of this shell-rich unit may be genetically unrelated, it is possible that the unit represents a single sediment gravity flow that underwent flow transformation, the upper layer representing deposition from a surging, turbulent flow that developed at the debris flow – seawater interface (hampton 1972; krause & oldershaw 1979; sohn 2000). swelling clay the facies consists of greasy, grey clay laminae or beds, 0.3–8 cm thick, which are recognised at four levels (83.15 m, 82.6 m, 78.95 m, 25.75 m). on contact with water, the clay shows prominent swelling (fig. 6b), suggestive of a significant smectite content within the clay sediment ( jeans et al. 2000). these thin swelling clay units are readily differentiated from the remainder of the mudstones in the core and may represent ashfall deposits that settled out from suspension, as reported from the jurassic of the north sea area ( jeans et al. 2000). they are thus identified as bentonites on the sedimentary log (fig. 2). geochemistry the average bulk geochemistry of the main facies is presented in tables 1 and 2. the sandstones show prominent variations in composition reflecting varying defacing page: fig. 8. representative photomicrographs (plane-polarised light) of the upper jurassic katedralen member (hareelv formation) in the blokelv-1 core; scale is applicable to all photographs. a: fine-grained sandstone (heterolithic facies) with abundant mudstone intraclasts (increasing in size and proportion upwards) sharply overlain by very fine-grained, quartz-dominated sandstone referred to the massive sandstone facies (14.55 m). b: very fine-grained sandstone (injectite; chaotic sandstone facies) with a high percentage of intergranular porosity shown by blue staining (94.20 m). c: fine-grained layered injectite (chaotic sandstone facies) with abundant sand-grade mudstone clasts in the upper part (61.96 m). d: fineto medium-grained, moderately sorted sandstone (massive sandstone facies) showing a high intergranular porosity (>25%) picked out by the blue stain (130.30 m). e: fine-grained sandstone injectite (chaotic sandstone facies) with large rip-out mudstone clasts probably incorporated during injection (102.39 m). f: muddy shell bed (bioclastic mudstone/sandstone facies); note the open clast-supported framework created by the bioclasts and the ubiquitous mud matrix (83.50 m). ctotal % stotal % 3.05 0.36 0.39 0.19 2.67 1.13 7.48 3.47 no. of samples sio2 % al2o3 % fe2o3 % mgo % cao % na2o % k2o % tio2 % p2o5 % mno % cr2o3 % loi* % sandstone (cemented) 10 64.89 7.94 3.97 3.02 5.89 0.71 table 1. average content of major elements in ankerite-cemented sandstones, sandstones, heteroliths and mudstones in the blokelv-1 core 1.86 0.41 0.10 0.08 0.006 10.9 sandstone 17 85.45 6.95 1.16 0.37 0.46 0.56 1.77 0.35 0.06 0.01 0.004 2.7 heterolith 5 68.29 12.72 4.13 0.98 0.97 0.71 2.46 0.62 0.13 0.03 0.010 8.7 mudstone * loss on ignition. 9 51.19 18.57 6.49 1.05 0.60 0.60 2.83 0.84 0.18 0.02 0.017 17.3 mbj art1 tabel 1 5454 40 50 60 70 80 90 100 sio 2 ( w t% ) 0 5 10 15 20 25 0 50 100 150 200 al 2o 3 ( w t% ) 0.2 0.4 0.6 0.8 1 1.2 n a 2 o (w t% ) 1 1.5 2 2.5 3 3.5 k 2 o (w t% ) 0 2 4 6 8 10 fe 2o 3 ( w t% ) 0 1 2 3 4 5 6 s t ot al (w t% ) 0 50 100 150 200 0 50 100 150 2000 50 100 150 200 0 50 100 150 2000 50 100 150 200 a depth (m)depth (m) gø02_13_121_mbj_blokelv.eps mbj art1 �g9afig. 9. geochemistry (selected major and trace elements) of the dominant facies of the blokelv-1 core, plotted against sample depth. samples with significant ankerite cement (where ca > 2 wt%) are differentiated from weakly or uncemented sandstones, heteroliths and mudstones; the illustrated linear trend lines are based only on the latter (non-ankeritic) population. ctotal and stotal refer to total weight percent of carbon and sulfur, respectively; loi, loss on ignition. 55 0 2 4 6 8 10 12 c ao (w t% ) 0 1 2 3 4 5 6 0 2 4 6 8 10 12 c to ta l ( w t% ) 0 5 10 15 20 25 lo i ( w t% ) 0 4 8 12 16 20 u (p pm ) 10 0 20 30 40 50 60 70 80 90 m o (p pm ) 0 50 100 150 2000 50 100 150 200 0 50 100 150 2000 50 100 150 200 0 50 100 150 2000 50 100 150 200 m go (w t% ) b depth (m)depth (m) sandstone heterolith mudstone trendlines turbidite, homogeneous injectite, homogeneous injectite with mudstone clasts turbidite with mudstone clasts sandstone-mudstone laminae with sandstone laminae laminated mudstone sandstone (uncemented) heterolith (uncemented) mudstone ankerite cemented gø02_13_122_mbj_blokelv mbj art1 �g9b 5656 grees of cementation in the samples (olivarius et al. 2018a, this volume). the geochemistry of the cemented intervals shows that the cement consists of ankerite (ca(fe,mg,mn)(co3)2). sandstones with ankerite cementation show the following parameters: ca > 2, mg > 1 and c > 1 wt% (fig. 9). about a third (10 from 29 samples) of the geochemically analysed sandstones are cemented with ankerite; the cement occurs in both the massive sandstone and chaotic sandstone facies, and one of the heterolith samples is also ankerite-cemented. selected geochemical depth-trends for the sandstones, heteroliths and mudstones are shown in fig. 9; the ankerite-cemented samples are excluded from calculations of linear trend-lines because they are not representative of the general sandstone population since ankerite formed preferentially in the finest-grained sandstones, possibly due to the presence of biogenic carbonate mud in this lithology (olivarius et al. 2018a, this volume). predictably, the ankerite-cemented sandstones generally show higher contents of fe, mg, mn, c and loi than the noncemented sandstones, which, in contrast, contain higher amounts of si (fig. 9, table 1). the clay content is clearly reflected in the bulk geochemistry: the average si-content in the non-cemented samples decreases from 85 wt% in sandstones to 68 wt% in heteroliths and to 51 wt% in mudstones, whereas the content of most other elements increases – aluminium, for example, with 7, 13 and 19 wt%, respectively (table 1). the si-content increases slightly up-section in the sandstones and heteroliths, whereas the amounts of al and k decrease slightly (fig. 9a). the na-content increases significantly upwards in the sandstones and heteroliths (from 0.3 to 0.9 wt% and from 0.5 to 1.1 wt%, respectively); no significant trend is recognised in the mudstones. the amount of u in the mudstones (fig. 9b) increases steadily and significantly up-section (from 3.7 to 17.5 ppm), whereas fe and s only show a minor increase in content up-section (fig. 9a). the three uppermost mudstone samples (i.e. in the interval 75–27 m) contain high levels of trace metals such as mo, u, co, ni, cu, zn and cd (fig. 9, tables 1, 2). the mudstones contain more al, fe, ti and p than recorded in average global mudstone compositions of various ages and areas (gromet et al. 1984) and less ca, mg, na and mn, when subtracting the loi values from the results. the high si-content in the sandstones reflects their quartz-rich nature (subarkoses), and the marked upwards increase in the na content in the sandstones and heteroliths points to a marked relative increase in albite grains within the feldspar population (fig. 9a). the high loi of 17.3 wt% in the mudstones (fig. 9b, table 1) results from a high organic content, since the loi measures the total content of structural water in both clay minerals and organic matter. the upwards increase in the content of fe and s in the mudstones is interpreted to correlate with an upward increase in the pyrite content, suggesting that most s is inorganic (fig. 9). redox-sensitive trace metals sandstone (cemented)  sandstone table 2. average content of minor and trace elements in ankerite-cemented sandstones, sandstones, heteroliths and mudstones in the blokelv-1 core heterolith mudstone no. of samples 10 17 5 9 ba 427 362 558 603 co 17.9 28.6 19.8 19.8 cs 1.8 1.3 2.9 7.6 ga 10.2 8.3 15.6 24.4 hf 6.5 5.6 8.9 5.8 nb 8.9 7.7 12.3 17.6 rb 65.7 56.6 85.4 125.2 sn 2 2 2 3 sr 170.9 84.6 131.4 170.1 ta 1.3 2.0 1.4 1.4 th 5.6 4.8 9.4 18.1 u 1.7 1.5 3.5 9.7 v 51 40 94 323 w 136.2 323.2 123.7 39.2 zr 228.4 204.2 316.2 208.1 y 16.9 12.3 22.5 41.5 la 19.8 16.4 29.6 47.8 ce 40.1 34.8 61.6 109.1 pr 4.88 4.09 7.34 12.34 nd 18.7 15.4 27.5 47.4 sm 3.27 2.66 4.95 8.65 eu 0.66 0.52 0.94 1.83 gd 2.87 2.27 4.27 7.63 tb 0.45 0.36 0.68 1.21 dy 2.48 1.97 3.69 6.65 ho 0.50 0.39 0.72 1.30 er 1.54 1.21 2.15 3.82 tm 0.24 0.19 0.34 0.59 yb 1.53 1.23 2.14 3.60 lu 0.23 0.19 0.33 0.53 mo 0.4 0.3 4.0 21.2 cu 5.1 10.3 14.0 66.9 pb 4.5 4.4 8.9 27.0 sc 8 5 10 17 zn 42 27 105 288 ni 11.2 6.5 33.8 95.0 as 3.0 0.9 7.0 37.7 cd 0.1 0.0 0.7 3.3 sb 0.0 0.0 0.1 0.7 bi 0.0 0.0 0.1 0.4 ag 0.0 0.0 0.1 0.6 au 0.6 1.5 0.0 0.1 hg 0.02 0.01 0.04 0.22 tl 0.0 0.0 0.1 0.1 se 0.3 0.1 1.9 13.6 mbj art1 tabel 2 all values are ppm, excepting au (ppb). values below detection limits are given as zero. 57 are increasingly insoluble under oxygen-depleted conditions so they are retained in the sediment (tribovillard et al. 2006). thus, the high content of trace metals such as mo, u and co in the upper part of the core supports sedimentological and organic geochemical evidence of increasing oxygen depletion and a high degree of preservation of marine organic matter in the upper kimmeridgian part of the core (see bojesen-koefoed et al. 2018, this volume). similarly, ni, cu, zn and cd are enriched in the upper part of the core; these elements were probably introduced into the system in organic matter and were preserved in association with the formation of pyrite (tribovillard et al. 2006). the mudstones have the highest rare-earth element (ree) concentrations, followed by the heteroliths and sandstones (fig. 10). all samples show a negative euanomaly, are enriched in the light-ree and have a flat heavy-ree chondrite-normalised pattern, corresponding to a typical crustal composition. the ratios between the ree are very similar for all samples, with the exception of the uppermost sandstone sample at 3.60 m, which has probably been affected by surface weathering. petrophysical core logs the petrophysical signatures of the blokelv-1 core are evaluated from the filtered bulk density log and spectral gamma-log traces in relation to stratigraphy and sedimentary facies (fig. 2). bulk density log sandstones are characterised by alternating units of low (2.35–2.5 g/cm3), medium (2.5–2.7 g/cm3) and high (2.7-2.9 g/cm3) filtered bulk densities (fig. 2). low bulkdensity signatures in sandstones are in general associated with porous zones (15–26%) in the massive turbidite sandstones and injectites (fig. 2; see also fig. 13). medium and high values are recorded in all described sandstone facies types and reflect differential diagenetic influence, particularly the presence/absence of ankerite, as discussed by olivarius et al. (2018a, this volume). mudstones are characterised by medium density values with a characteristic low-density interval from 163–153 m. the four igneous intrusions are characterised by very highdensity readings (c. 3 g/cm3). spectral gamma log the total gamma log (grtotal ) shows distinct log patterns that can be related to the sedimentary facies (fig. 2). massive sandstone units, both high-density turbidites and injectites, typically display a box-shape log motif with low gr values (c. 50 api) due to a low content of radioactive minerals. sandstones with abundant mudstone intraclasts show slightly higher gr values. the potassium and thorium values locally show upward-increasing trends in individual sandstone units (e.g. 142.7–134.7 m). the heterolithic mudstone–sandstone facies shows intermediate gr values. intervals with increasing mudstone/sandstone ratios up section display clear upward-increasing trends in u, k and th values (e.g. 188.7–186.5 m). the mudstones are characterised by grtotal values of 150–200 api, with an overall increasing upward trend (fig. 2). three main intervals are distinguished based on the gr trends in the mudstones of the cored section, primarily resulting from the uranium content; potassium and thorium show no overall stratigraphic trends. the lower interval (233–185 m), assigned to the middle oxfordian – lower upper oxfordian (alsen & piasecki 2018, this volume), is characterised by 5–10 ppm uranium concentrations that up-section define an increasing la ce pr nd sm eu gd tb rare-earth elements dy ho er tm yb lu re e sa m pl e / r ee c ho nd rit e sandstone heterolith mudstone turbidite, homogeneous injectite, homogeneous injectite with mudstone clasts turbidite with mudstone clasts sand/mudstone laminae with sandstone laminae laminated mudstone 1 10 100 1000 gø02_13_123_mbj_blokelv mbj art1 �g10 fig. 10. rare-earth elements (ree) diagram showing highest ree concentrations in the mudstones, intermediate ree concentrations in the heteroliths and lowest ree concentrations in the sandstones. the ree abundances are normalised to the chondrite values of anders & grevesse (1989), multiplied by 1.36 as proposed by korotev (1996), to obtain ree estimates close to the early ordinary chondrite compositions and to ensure consistency with older literature. 5858 trend followed by a decreasing trend. the middle succession (185–103 m), assigned to the lower upper oxfordian – lower kimmeridgian (alsen & piasecki 2018, this volume), is characterised by 8–12 ppm uranium concentrations that depict an upward-increasing trend in the lower part (178–168 m) and two upward-decreasing trends from 128–119 m and 110–103 m. the mudstone units in the upper succession (103–10 m), assigned to the lower–upper kimmeridgian (alsen & piasecki 2018, this volume), are characterised by 10–15 ppm uranium concentrations, and show a prominent upward increasing trend from 91–83.5 m. the mudstone units in the upper levels of this interval (23–10 m) show a progressive upward decrease in the grtotal and uranium values. the total gamma values of the igneous intrusions are very low (<10 api) and are associated with a very low potassium content of <0.3% (fig. 2); detailed geochemistry of the intrusions is given in larsen (2018, this volume). discussion processes and palaeoenvironment deposition of the blokelv-1 cored succession occurred almost exclusively under severely oxygen-depleted seafloor conditions in a deep-water marine setting at the transition from lower slope to basin floor. background sedimentation was of dark grey to black laminated mud comparable to that described from the hareelv formation at outcrop (surlyk 1987). the bioturbated mudstone facies that is restricted to the lower part of the core reflects a dysoxic environment; the limited nature of this occurrence, however, precludes interpretation of the spatial and temporal significance of this facies. massive fineto medium-grained sand beds were deposited from sediment gravity flows, probably in confined slope gullies and/or as unconfined sheets, lobes or splays on the basin floor (surlyk et al. 2007). the dominant transport mechanism is interpreted to have been high-density turbidity currents that probably evolved from sandy debris flows initiated by episodic collapse of the sandy shelf margin (surlyk et al. 2007). it is likely that such collapse events at the margin were seismically triggered given the fault-associated differential subsidence of the westward tilted jameson land basin during the late jurassic proposed by bjerager et al. (2018b, this volume). the deposited sand was prone to post-burial remobilisation by fluidisation and injection to form geometrically complex bodies. this probably resulted from cyclic loading triggered by seismic shocks associated with the rift climax (surlyk & noe-nygaard 2001; surlyk et al. 2007); a similar trigger mechanism has been suggested for palaeogene remobilised, injected deep-water sandstone in the north sea region (e.g. duranti & hurst 2004; hamberg et al. 2005). some small-scale sandstone injectites (e.g. dykes) that are demonstrably folded were formed during early burial and subsequently folded during continued burial and compaction. post-burial remobilisation of sandstones occurred to varying degrees and transitions are observed from sandstones with local water-escape structures to completely remobilised structureless sandstones. sandstone injectites commonly include isolated mudstone clasts that were ripped from the host rock during injection; in places these form mudstone-clast breccias with a sandy matrix (surlyk et al. 2007). the remobilised sandstone units recognised in the blokelv-1 core (fig. 2) amount to about two-thirds of the total sandstone units, and it is considered likely that this is a conservative estimate. the sandstones were sourced from the west or north, as indicated by their zircon u–pb ages and heavy mineral assemblages (olivarius et al. 2018b, this volume). the same broad sediment transport trends have been reported from sedimentological data from oxfordian shelf-edge deltas of the olympen formation (larsen & surlyk 2003; bruhn & surlyk 2004) and from sandy shelf-margin clinoforms of the volgian raukelv formation (surlyk & noe-nygaard 1991, 2005). the heterolithic intervals are interpreted to have been deposited mainly from low-density turbidity currents. the common evidence of synand post-depositional deformation in this facies, such as convolute and distorted bedding and slumping, is suggestive of a high sediment influx in the lower slope – basin floor setting. the intervals dominated by heterolithic sediments may have been associated with peaks in progradation of sandy shelf deposits, represented in part by the olympen and raukelv formations. the heteroliths may thus be considered, at least in part, as prodeltaic hyperpycnites, analogous to comparable facies described from cretaceous basins in north america (e.g. soyinka & slatt 2008; bhattacharya & maceachern 2009). alternatively, this rhythmic heterolithic facies can be compared with levee deposits described from turbidite channel complexes (e.g. posamentier & walker 2006; butterworth & macdonald 2007; kane et al. 2007), although such channel systems have not to date been identified in the exposed sections of the jameson land basin. previous outcrop-based studies have interpreted the massive sandstone bodies to represent line-sourced systems that lacked discrete fan development (surlyk 2003; surlyk et al. 2007). 59 basin evolution: implications of blokelv-1 the tectonostratigraphy and relative sea-level history of the jameson basin was presented by surlyk (2003). the late oxfordian – early volgian period was characterised by increasing rifting and marine flooding in the region; sequence stratigraphic analysis based on the succession in milne land (west of jameson land) suggested maximum flooding in the late kimmeridgian (a. eudoxus chron) to early volgian (p. elegans chron; surlyk 2003). the coarse-grained shelf margin wedge of the raukelv formation, exposed in southern jameson land (fig. 1), developed during a long-term sea level fall that started in the early volgian and culminated in late volgian time (surlyk & noe-nygaard 2005). the blokelv-1 core provides a complete middle oxfordian – earliest volgian record from the central part of the jameson land basin (alsen & piasecki 2018, this volume). the depositional evolution of the cored interval is presented in a regional and basinal context in a separate paper (bjerager et al. 2018b, this volume). long-term depositional trends are clearly recorded in the grtotal and spectral gamma (uranium) data from the mudstone intervals in the blokelv-1 core. the mudstones show an overall upward increase in grtotal and uranium values from the middle oxfordian to the lower kimmeridgian, culminating in the middle–upper kimmeridgian, followed by a decrease in grtotal and uranium values in the uppermost kimmeridgian – lower volgian (fig. 2, bjerager et al. 2018b, this volume). these trends correlate with an overall shift in the composition of the organic material in the mudstones, which is predominantly marine but has a high terrestrial proportion in the deeper part of the core (bojesen-koefoed et al. 2018, this volume). the terrestrial component of the organic matter decreases upward in the core in the interval 233.8–100 m (bojesen-koefoed et al. 2018, this volume); this change in the balance between terrestrial and marine organic input may reflect the overall major transgression that culminated in the late kimmeridgian in the jameson land basin (surlyk 2003). the parallel long-term increase in the grtotal and uranium values, in association with other geochemical parameters (mo, stotal), reflects increasing oxygen depletion on the sea floor. this was probably also linked to the regional transgressive trend discussed above, though basinal isolation and inhibition of circulation may also have contributed to sea-floor anoxia, particularly in the later regressive phase (early volgian). minor, shorter-term depositional trends are also recorded in the blokelv-1 core. they are reflected both by vertical stacking of facies and by variations in grtotal and uranium values in mudstone units (fig. 2). upward increases in uranium concentrations in the mudstone units probably reflect increasing organic content; this may have resulted from increasingly oxygen-deficient conditions at the seafloor favouring preservation of organic matter and/or a relative decrease in siliciclastic input relative to organic matter. upward-decreasing uranium concentrations in mudstones can be commonly correlated with increasing siliciclastic input, as indicated by an increase in siltstone and fine-grained sandstone laminae; in some cases, such coarsening-upward trends culminate in heterolithic intervals and massive sandstones (fig. 2). the mediumto coarse-grained sandstones of the sjællandselv member at the top of the cored section represent amalgamated gravity-flow sandstones that are inferred to have resulted from collapse of the prograding shelfmargin sandstones of the raukelv formation, recording a significant relative lowering of sea level (surlyk & noenygaard 2005). reservoir potential the porosity and permeability of the sandstones and heteroliths, as measured by conventional core analysis, show good correlation and the measured values are compatible with the log-derived porosity/permeability relationship (fig. 11). sandstone bodies with porosities >15% are recorded throughout the cored succession, but sandstones with porosities >20% are restricted to the upper 130 m of the core. sandstone beds or units with low to medium porosities (<15%) occur throughout the section. 0 5 10 15 20 25 30 porosity (%) 0.01 0.001 0.10 1 10 100 1000 g as p er m ea bi lit y (m d ) log-derived data high-density turbidite (massive sandstone) injectite (chaotic sandstone, mudstone-clast breccia) low-density turbidite (heterolith) gø02_13_140_mbj_blokelv mbj art1 �g11 fig. 11. porosity and permeability cross-plot of the three main sandrich facies based on conventional core analysis and compared with log-derived values. 6060 the massive high-density turbidite sandstones and the injected chaotic sandstones have net thicknesses of 35 m and 48 m, respectively (fig. 2). they share similar reservoir properties with porosities of 14–26% for c. 90% of the analysed samples, and air-permeabilities of 2–400 md in 80% of the samples. the remaining samples range in porosity between 7 and 10% and in air-permeability between 0.05 and 0.5 md. the log-derived porosities of the two facies types correlate well with the measured values although the log-derived porosity values are slightly lower in the most porous beds (fig. 12). sandstone units over 1 m in thickness with porosities over 15% have a net thickness of about 40 m in the blokelv-1 core (fig. 12), giving a net/gross ratio for the entire section of 17%. the grain density of the sandstones is in the range of 2.62– 2.75 g/cm3, with an average of 2.68 g/cm3. high porosity (>15%) sandstones show densities of 2.62–2.69 g/cm3, whereas low porosity (<10%) sandstones have densities of 2.75 g/cm3 (fig. 12). the heterolithic sandstone–mudstone samples have porosities of 2–14% and permeabilities of 0.1–0.6 md (figs 11, 12); the corresponding grain densities are in the range of 2.65–2.75 g/cm3 (fig. 13). porosity in mudstones is in the range of 1.6–5.6% (fig. 11); permeability is below the detection limit (0.05 md). mudstone grain density ranges between 2.37 and 2.57 g/cm3 (fig. 13). the lowest value in mudstones was recorded in the interval 163–153 m (fig. 2), possibly reflecting abundant coal clasts or other organic particles that may have grain densities as low as 1 g/cm3; this is compatible with total organic carbon (toc) analytical data in bojesenkoefoed et al. (2018, this volume) showing peak values in mudstones in the interval 170–140 m. igneous intrugø04_02m_453_03_mbj_blokelv_sed mbj fig 12 permeability (md) high-density turbidite sandstones injectite sandstone heterolithic sandstone/mudstone igneous intrusion mudstone analytical data porosity % 10 depth (m) 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 210 220 230 10 depth (m) 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 210 220 230 233.8 be be clay si sand pebbl. 0.1 10 0 10 20 mbj art1 �g12 gø04_02m_453_03_mbj_blokelv_sed mbj fig 12 permeability (md) high-density turbidite sandstones injectite sandstone heterolithic sandstone/mudstone igneous intrusion mudstone analytical data porosity % 10 depth (m) 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 210 220 230 10 depth (m) 20 30 40 50 60 70 80 90 100 110 120 130 140 150 160 170 180 190 200 210 220 230 233.8 be be clay si sand pebbl. 0.1 10 0 10 20 mbj art1 �g12 fig. 12. log-derived porosity and permeability data plotted against depth and the generalised sedimentary log. for comparison, the measured porosity and permeability values from conventional core analysis of the main facies (see fig. 11) are also shown. for legend, see fig. 2. 61 sions have porosities <2% and permeabilities <0.05 md (fig. 12); grain densities range between 3.00 and 3.05 g/ cc (figs 2, 13). ankerite cementation has had a marked influence on the reservoir properties of the blokelv-1 core sandstones; most sandstone plugs with low porosity and low permeability have a high degree of ankerite cement (olivarius et al. 2018a, this volume). in general, the grain densities of the sandstones reflect the degree of cementation with ankerite such that sandstones with no cement have grain densities close to 2.65 g/cm3 and heavily cemented sandstones have grain densities of c. 2.75 g/cm3. sandstones with high densities (c. 2.68–2.76 g/cm3) are recorded, for example, at 151.65 m and 166.70 m (fig. 2); such occurrences are in general attributed to ankerite cement, which has a grain density of 3.00–3.11 g/cm3. calculation of the average grain density of a typical sandstone with 30 wt% ankerite and 70 wt% quartz results in an average grain density of 2.75 g/cm3. one sandstone sample from the blokelv-1 core studied by olivarius et al. (2018a, this volume) showed an ankerite content of about 41 wt%, resulting in significant reduction in the porosity and permeability. sandstones in the vicinity of the igneous intrusions (e.g. at 56.55 m) also show low porosities and permeabilities and high grain densities attributable to ankerite precipitation in the pore spaces. both the bulk density log and the bulk geochemistry can thus be used to identify the intervals with ankerite cementation. the sandstone units showing high bulk densities (>2.7 g/ cm3) are 0.25–2.95 m thick and occur throughout the cored interval; they amount to a combined thickness of 6.6 m that is considered to represent the net thickness of sandstone affected by pervasive ankerite cementation. spherical concretions commonly occur in massive sandstone units in exposed sections of the hareelv formation (surlyk & noe-nygaard 2001); it is considered most likely that these structures are similarly cemented by ankerite. if the pervasively ankerite-cemented units recorded in the blokelv-1 core show comparable morphologies to the concretions at outcrop, i.e. have similar limited lateral continuity, then they will have little influence on overall properties at reservoir scale. summary and conclusions deposition of the blokelv-1 cored section occurred in relatively deep water in the central part of the jameson land basin in lower-slope to basin-floor environments, in accordance with published outcrop studies in the area (e.g. surlyk 1987). parallel-laminated black organic-rich mudstones were deposited primarily by hemipelagic suspension under anoxic conditions; intercalated rare to abundant laminae of siltand fine sandstone were deposited from low-density turbidity currents. laminated hete rolithic sandstone–mudstone intervals were deposited from low-density turbidity currents and fall-out from suspension. massive sandstones are interpreted to have been deposited from high-density turbidity currents and were susceptible to post-burial remobilisation and injection into the surrounding host mudrock. the resulting succession comprises nearly equal amounts of intercalated mudstone and sandstones in geometrically complex bodies and thus confirms published studies on outcrop sections (e.g. surlyk et al. 2007). at a conservative estimate, about two-thirds of the sandstone units were remobilised to some degree after burial. the facies analysis, bulk geochemistry and high-resolution spectral gamma log data facilitate recognition of long-term and superimposed short-term depositional trends in the continuous middle oxfordian – lowermost volgian succession provided by the blokelv-1 core. the spectral gamma log measured on core indicates a positive 0 5 10 15 20 25 30 2.30 2.40 2.50 2.60 2.70 2.80 2.90 3.00 3.10 high-density turbidite injectite heterolith mudstone basalt g ra in d en sit y (g /m l) porosity (%) gø02_13_141_mbj_blokelv mbj art1 �g13 fig. 13. cross-plot of grain density and porosity of the main facies based on conventional core analysis. the mudstones show low grain density and low porosity; igneous intrusive rocks show low porosity and high grain density. the variability within the heterolith facies (medium–low porosity, intermediate but variable grain density) reflects the lithological heterogeneity in the facies. high-density turbidite sandstones and injectites lacking pervasive ankerite cement show mainly high porosity (>15%) and relatively constant intermediate grain density. most of the sandstones with low porosity (<10%) show high grain density reflecting the ankerite content (see text for discussion). 6262 correlation between the uranium concentration and marine-derived organic material; the high values of certain trace metals in the upper part of the core are indicative of severe oxygen depletion on the sea floor. the methodology utilised in this study to derive reservoir properties (porosity/permeability) from the core density log and core spectral gamma log traces provides a highly reliable dataset that is compatible with the conventional core analysis data. the resulting quantitative and qualitative data sets are useful in evaluating the reservoir quality of the upper jurassic deep-water sandstones of the jameson land basin. the high-density turbidite sandstones and sandstone injectites typically show comparable, good reservoir properties with porosities of 15–26% porosity and permeabilities of 2–200 md. the sandstone intervals showing such reservoir potential amount to a net vertical thickness of about 40 m, corresponding to a net/gross ratio of 17% for the entire cored interval. in contrast, the sandstone–mudstone heteroliths show poor reservoir quality, typically with <14% porosity and <0.6 md airpermeability. the massive sandstones thus typically have high porosities and show potential as reservoir units. the reservoir volume depends, however, on their thickness and lateral extent. the sandstones are interpreted to have been deposited on the slope, filling gullies that transected the slope, or as laterally extensive sheets on the lower slope and basin floor (surlyk & noe-nygaard 2001; surlyk et al. 2007). post-depositional remobilisation and injection of the sands resulted in a complex 3d distribution of potential reservoir units within the mudstone succession. although this complexity reduces the architectural predictability of the potential reservoir, the injection process probably created significant vertical connectivity between individual sandstone bodies, adding to the effective connected reservoir volume. factors that reduce the reservoir quality of the sandstones in the blokelv-1 core include the abundant sandgrade mud clasts, ankerite cements and cements associated with igneous intrusions. variation in the porosity and permeability of ankerite-free sandstones is controlled by a range of diagenetic processes – both negative, such as diagenetic clay minerals, and positive, such as the precipitation of microquartz coatings that inhibit porosityoccluding macroquartz cementation (olivarius et al. 2018a, this volume). acknowledgements we thank john boserup, jette halskov and stefan sølberg for sampling assistance and figure preparation. hans jørgen lorentzen and marga jørgensen are acknowledged for the ccal analytical work. the referees, michael larsen and finn surlyk, provided pertinent and constructive comments that improved the paper significantly. references alsen, p. & piasecki, s. 2018: biostratigraphy of the hareelv formation (upper jurassic) in the blokelv-1 core, jameson land, central east greenland. in: ineson, j. & bojesen-koefoed, j.a. 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(eds): sand injectites: implications for hydrocarbon exploration and production. aapg memoir 87, 141–149. tribovillard, n., algeo, t.j., lyons, t. & riboulleau, a., 2006. trace metals as paleoredox and paleoproductivity proxies: an update. chemical geology 232, 12–32. _________________________________________________________________________________________ manuscript received 9 november 2015; revision accepted 1 may 2018 geological survey of denmark and greenland bulletin 6, 5-28 5 lower palaeozoic stratigraphy of the east greenland caledonides m. paul smith, jan audun rasmussen, steve robertson, a.k. higgins and a. graham leslie the lower palaeozoic stratigraphy of the east greenland caledonides, from the fjord region of north-east greenland northwards to kronprins christian land, is reviewed and a number of new lithostratigraphical units are proposed. the slottet formation (new) is a lower cambrian quartzite unit, containing skolithos burrows, that is present in the målebjerg and eleonore sø tectonic windows, in the nunatak region of north-east greenland. the unit is the source of common and often-reported glacial erratic boulders containing skolithos that are distributed throughout the fjord region. the målebjerg formation (new) overlies the slottet formation in the tectonic windows, and comprises limestones and dolostones of assumed cambrian–ordovician age. the lower palaeozoic succession of the fjord region of east greenland (dominantly limestones and dolostones) is formally placed in the kong oscar fjord group (new). amendments are proposed for several existing units in the kronprins christian land and lambert land areas, where they occur in autochthonous, parautochthonous and allochthonous settings. keywords: early palaeozoic, north-east greenland, stratigraphy. m.p.s., lapworth museum, school of geography, earth and environmental sciences, university of birmingham, edgbaston, birmingham b15 2tt, uk. e-mail: m.p.smith@bham.ac.uk j.a.r. & a.k.h., geological survey of denmark and greenland, øster voldgade 10, dk-1350 copenhagen k, denmark. s.r. (deceased) & a.g.l., british geological survey, murchison house, west mains road, edinburgh eh9 3la, uk. the east greenland caledonides extend for over 1300 km between scoresby sund (70°n) and kronprins christian land (81°30′n), cropping out in an ice-free coastal strip of variable width (fig. 1). in the south, the exposed width of the orogen is 300 km, but in the north this is reduced to less than 100 km. the orogen has been the subject of a series of systematic mapping programmes by the survey since 1968, supplementing and revising work by other groups, most notably lauge koch’s long series of geological expeditions between 1926 and 1958 (see haller 1971). this paper documents the key stratigraphical observations on lower palaeozoic rocks made during expeditions from 1993 to 1995 by the former geological survey of greenland (ggu) to kronprins christian land and lambert land (78°–82°n), and the 1997–1998 field work by the survey in the kong oscar fjord region (72°–75°n). a number of new lower palaeozoic lithostratigraphical units are proposed on the basis of this work and amendments are proposed for some existing units. the lower palaeozoic sediments of kronprins christian land and lambert land (fig. 2) were deposited on a subtidal to peritidal platform that constituted the south-easternmost part of the franklinian basin (higgins et al. 1991), and lay at a marked inflexion of the laurentian margin where it turns through 90° between the present-day e–w-trending coast of north greenland and the n–s-trending coast of east greenland. the stratigraphy erected in the caledonian foreland of eastern north greenland (peary land and western kronprins christian land; peel 1985; higgins geological survey of denmark and greenland bulletin 6, 5–28 © geus, 2004 geus bulletin 6.pmd 10-02-2005, 09:535 6 ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ c a le d o n id e s greenland 70°n 74°n 78°n 82°w 35°w wandel sea centrumsø 100 km jameson land scoresby sund 25°w traill ø wollaston forland bessel fjord danmarkshavn lambert land peary land station nord kr on pr ins c hr ist ian l an d nørreland window kronprins christian land thin-skinned thrust belt (parautochthonous foreland) palaeogene basalts palaeogene intrusions wandel sea basin: carboniferous–palaeogene sediments east greenland basins: carboniferous–cretaceous sediments devonian – continental sediments late to post-kinematic granites neoproterozoic–ordovician sediments (east greenland) neoproterozoic–silurian sediments (eastern north greenland) palaeo-mesoproterozoic sediments and basalts (eastern north greenland) crystalline complexes and sediments (archaean–mesoproterozoic) neoproterozoic–silurian sediments (north greenland) palaeo-mesoproterozoic sediments and basalts (eastern north greenland) mainly crystalline rocks – parautochthonous windows thrust fault/shear zone tectonic zone boundary post-caledonian late to post-caledonian caledonian orogenic belt caledonian foreland ▲ ▲ fig. 8 fig. 2 c a le d o n ia n so le t h ru st in la n d ic e m a r g in a l t h r u s t b e lt t h ic k s k in n e d t h r u s t b e lt gåseland window charcot land window eleonore sø window målebjerg window hamberg gletscher foreland dronning louise land ▲ ▲ ▲ fig. 1. geological map of the east greenland caledonides, showing location of the foreland windows in the western marginal thrust belt. modified from higgins & leslie (2000). frames indicate the regions shown at larger scales in figs 2 and 8. geus bulletin 6.pmd 10-02-2005, 09:536 7 et al. 1991; smith & bjerreskov 1994) can, to a significant degree, be applied to the successions farther to the east and south within the study area (figs 2, 3), although many of the units differ in detail. this region is the only part of the franklinian basin to have been subsequently affected by the caledonian orogeny, and the lower palaeozoic units described here occur both in the foreland and in the thin-skinned parautochthonous fold-and-thrust belt beneath the vandredalen thrust sheet (fig. 2; higgins et al. 2001b). the lower palaeozoic successions of the fjord region of the southern parts of north-east greenland (72°–74°30′n; see fig. 8) have been the subject of a more protracted research effort that extends back to the 19th century. the cambrian–ordovician of the fjord region has become one of the classic reference areas for this stratigraphical interval, but it was only during the 1997–1998 survey field seasons that the tectonic context was fully elucidated (higgins et al. 2004a). it is now clear that the cambrian–ordovician of the fjord region, together with the underlying, neoproterozoic tillite group and eleonore bay supergroup, make up the upper part of the highest thrust sheet in the orogen, which has been transported several hundred kilometres from the east-south-east (higgins & leslie 2000; higgins et al. 2004a). furthermore, notably different developments of lower palaeozoic rocks were discovered cropping out in tectonic windows adjacent to the inland ice; the successions within these windows are disturbed by caledonian deformation, and are interpreted as autochthonous to parautochthonous representatives of the foreland (smith & robertson 1999a, b; higgins et al. 2004a, b, this volume). the latter are fully documented for the first time here, and provide data critical for interpreting the lower palaeozoic evolution of the iapetus passive margin and its subsequent deformation during the caledonian orogeny. kronprins christian land the lower palaeozoic units cropping out over a c. 5000 km2 area in southern kronprins christian land, eastern north greenland (figs 2, 3), were mapped and documented in 1993–1995 as part of the 1:500 000 mapping programme carried out by ggu. within the thin-skinned parautochthonous fold-and-thrust belt forming much of this region (higgins et al. 2004b, this volume) exposure is often poor due to an extensive cover of recent glacio-fluvial sediments, and the best exposures are on the slopes of the main valleys, with more scattered outcrops on the plateaus. the discontinuous exposure is disrupted by a series of eastward-dipping thrusts with displacements of several hundred metres to a few kilometres. correlation of superficially similar middle ordovician to silurian sediments in alternating units of peritidal dolostone and subtidal burrow-mottled limestones required detailed sedimentary facies analysis. macrofaunal biostratigraphy in the field was supplemented by the biostratigraphic analysis of conodonts, which in almost all cases verified the field determinations (rasmussen & smith 2001). cambrian sediments are restricted to the sandstones of the kap holbæk formation (see below), scattered representatives of which occur in autochthonous, parautochthonous and allochthonous settings. kap holbæk formation revised history. the kap holbæk formation was first documented by adams & cowie (1953) during a geological reconnaissance around the head of danmark fjord, and was informally divided into five members. fränkl (1954) demonstrated that the unit was also present in kronprins christian land around sæfaxi elv, and hurst & mckerrow (1985) interpreted it as occurring within thrust sheets north of romer sø. initial biostratigraphical determinations (peel & vidal 1988) concluded that, on the basis of acritarchs, the unit was of latest vendian (late ediacaran) age. clemmensen & jepsen (1992) revised the hagen fjord group of haller (1961) to encompass a phase of neoproterozoic shallow marine sedimentation which succeeds the palaeoproterozoic–mesoproterozoic independence fjord group and the mesoproterozoic zigzag dal basalt formation, and pre-dates franklinian basin sedimentation. since the kap holbæk formation was thought to be of vendian age, the formation was considered to be representative of the youngest phase of sedimentation in the hagen fjord group basin. however, two schemes of possible stratigraphic correlation could still be used to express the relationship of the kap holbæk formation to the underlying fyns sø formation, with radically different consequences in terms of basin evolution models and correlation with east greenland. these alternatives were outlined by sønderholm & jepsen (1991): geus bulletin 6.pmd 10-02-2005, 09:537 8 ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲▲ ▲ ▲ ▲ ro mer sø i n go lf fjord amdrup land holm land hovgaard ø kap bernhoft dijmphna sund hek la sund centrumsø sk a l l in g e n syd vej dal r ivierada l 20°w 80°n 81°n blåsø sk jo ld un ge elv nioghalvfje rdsfjorden græ sel v sæ faxi elv d an m ar k fj or d va nd re d al en th ru st va nd re da le n kap holbæk p c m a hj lambert land m fl h wandel sea basin sequence (post-caledonian) samuelsen høj formation lauge koch land formation odins fjord formation turesø formation wandel valley formation kap holbæk formation crystalline basement thrust fault shear zone børglum river and sjælland fjelde formations fyns sø, kap bernhard, campanuladal fms hagen fjord gp rivieradal group independence fjord gp and basaltic formations vandredalen thrust hagen fjord group ▲ ▲ ▲ ▲ vandredalen thrust sheet ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲▲ ▲ ▲ c a le d o n id es greenland d p 25 km geus bulletin 6.pmd 10-02-2005, 09:538 9 1. the late ediacaran kap holbæk formation could rest conformably on the fyns sø formation, in which case the whole of the hagen fjord group would probably be of vendian age (sønderholm & jepsen 1991, fig. 18b). 2. the formation could rest disconformably on the fyns sø formation; the latter could then be representative of the late riphean to sturtian carbonate developments that characterise the north atlantic area. in this case, a substantial hiatus would span the vendian, and the hagen fjord group would be of riphean to sturtian age (sønderholm & jepsen 1991, fig. 18a). the kap holbæk formation overlies stromatolitic dolostones of the fyns sø formation at kap holbæk with a well-defined, but in detail obscured, contact. clues that might have helped to differentiate between these two hypotheses include the occurrence of a single clast of fyns sø formation-like lithology in the morænesø formation, a glacially influenced deposit of presumed varanger age in central peary land (collinson et al. 1989; sønderholm & jepsen 1991). an additional indication was provided by fränkl (1955) who had interpreted the kap holbæk formation as being the fill of karstic cavities in the top of the fyns sø formation at a single locality in sæfaxi elv. field work in sæfaxi elv in 1994 confirmed and expanded on fränkl’s observations. the uppermost part of the fyns sø formation was found to contain cave systems, infilled with sandstone, which are of typical phreatic character (smith et al. 1999). the fyns sø formation is unconformably overlain by the wandel valley formation, which is sandy at the base but passes upwards into typical dolomitic lithofacies of the danmark fjord member. confirmation of the lithostratigraphical relationships came from a locality at the junction of hjørnegletscher and ingolf fjord; here the fyns sø formation is overlain by the kap holbæk formation but infilled caves are preserved in the uppermost fyns sø formation (smith et al. 1999). deep channels are incised into the top of the latter and both the caves and the channels are infilled by lithofacies that can be directly matched to the lower part of the kap holbæk formation (smith et al. 1999). the channels and caves constitute one of the most spectacular and well-preserved examples of pre-carboniferous palaeokarst recorded to date. the demonstration of a hiatus between the fyns sø and kap holbæk formations lends support to correlation alternative 2 above (scheme a of sønderholm & jepsen 1991), which invokes a major time gap between the two units (fig. 3). further support comes from the interpretation of bjørnøya (western svalbard) as a detached north greenland terrane (smith 2000). on bjørnøya, a carbonate unit similar to the fyns sø formation is overlain by a sparse diamictite interpreted as a glacial horizon by harland et al. (1993), indicating that the carbonate is pre-vendian. the glacial unit, the sørhamna formation, is unconformably overlain by the ‘younger dolomite’, a peritidal dolostone containing a mid-early ordovician fauna that is closely comparable with the wandel valley formation (smith 2000). efforts to internationally standardise the position of the vendian–cambrian boundary have produced a wealth of biostratigraphic data in recent years. the redefined boundary has the effect of including a considerable time interval, the nemakit–daldyn (544–535 ma), within the early cambrian that was previously included within the neoproterozoic. furthermore, deep skolithos burrows are now known to have a first appearance in the tommotian (mid-early cambrian). their presence close to the base of the kap holbæk formation means that the unit is of early cambrian age, and is thus a probable correlative of the buen formation of north greenland, as originally postulated (peel 1980; peel et al. 1981; see fig. 7). because of the substantial time gap between the fyns sø formation and the kap holbæk formation, probably corresponding to the whole of the vendian, the younger unit is here removed from the hagen fjord group. type section. kap holbæk, a headland close to the head of danmark fjord on its western side (fig. 2). thickness. estimates of the thickness of the kap holbæk formation in the vicinity of danmark fjord range from 135 m (adams & cowie 1953) to 150 m (clemfacing page: fig. 2. geological map of kronprins christian land and lambert land; see inset outline map and fig. 1 for location. the region shows the well-exposed transition from autochthonous foreland in the west around kap holbæk and danmark fjord, eastwards through a thin-skinned fold-andthrust belt, to the allochthonous vandredalen thrust sheet and higher thrust sheets. fl, finderup land; h, harefjeld; hj, hjørnegletscher; m, marmorvigen; pcma, prinsesse caroline-mathilde alper. on inset map of greenland; d, daugaard-jensen land; p, peary land. geus bulletin 6.pmd 10-02-2005, 09:539 10 mensen & jepsen 1992). farther to the east, in the inner part of ingolf fjord, the formation is 180 m thick but thins southwards, along the western flanks of the prinsesse caroline-mathilde alper, to zero around sæfaxi elv. at sæfaxi elv, the ordovician wandel valley formation rests unconformably on the fyns sø formation and kap holbæk sediments are present only as the fill of palaeokarst cavities (smith et al. 1999). the allochthonous vandredalen thrust sheet does, however, contain the kap holbæk formation (see below), indicating that the formation was present farther to the east prior to thrusting (higgins et al. 2001b). the thinning of the formation to zero is coincident with the position of the rift shoulder of the hekla sund basin, the name given by fränkl (1955) to the half-graben rift basin in which the neoproterozoic rivieradal group accumulated (see also smith et al. 1999; higgins et al. 2001b; smith et al. 2004, this volume), suggesting that the structure continued to exert an influence on sedimentation into the earliest palaeozoic. boundaries. at the type locality, the kap holbæk formation disconformably overlies the fyns sø formasilurian ordovician cambrian vendian sturtian riphean lauge koch land formation samuelsen høj formation odins fjord formation turesø formation børglum river formation sjælland fjelde formation wandel valley formation kap holbæk formation hagen fjord group fyns sø fm kap bernhard fm campanuladal fm jyske ås fm zig-zag dal basalt formation independence fjord group hekla sund fm, aage berthelsen gletscher fm, & interbedded quartzites thermal subsidence extensional rifting and block tilting extensional riftingrivieradal group (allochthonous vandredalen thrust sheet only) post-rift thermal subsidence baltica collision thrust loaded flysch basin tectonic setting depositional environment stratigraphy lapetus passive margin lapetus opening pre-lapetus rift-sag cycle intracratonic extensional events thermal subsidence block tilting zz if hs/ab kh rg    fig. 3. pre-caledonian stratigraphy and tectonic history of kronprins christian land and lambert land. modified from smith et al. (1999). geus bulletin 6.pmd 10-02-2005, 09:5310 11 tion with a hiatus that probably corresponds to most of the vendian. it is disconformably overlain by the wandel valley formation, which is mid-early ordovician in age at the base (smith 1991; smith & bjerreskov 1994). these relationships pertain throughout the outcrop area, including the palaeokarst localities (smith et al. 1999). distribution. the kap holbæk formation crops out around the southern end of danmark fjord, extending southwards along the western side of skjoldungeelv to the inland ice (fig. 2). farther east, it crops out along the western side of ingolf fjord south of hjørnegletscher and along the western flanks of the prinsesse caroline-mathilde alper. at the junction of hjørnegletscher and ingolf fjord it additionally occurs infilling fossil caves in the fyns sø formation, and this is the only context in which it occurs around sæfaxi elv. the formation is also present in the frontal part of the vandredalen thrust sheet in finderup land, north of romer sø, where it was first described in the context of the ‘finderup land nappe’ (hurst & mckerrow 1981a, b; hurst et al. 1985). biota and age. peel & vidal (1988) noted a low diversity palynomorph assemblage from the formation and considered that the flora ‘in general suggests an age older than cambrian’. however, as noted above, the base of the cambrian has now been extended significantly downwards. the sandstones of the kap holbæk formation also contain deep skolithos burrows which extend vertically for many tens of centimetres (clemmensen & jepsen 1992, fig. 29) and occur to within 30 m of the base (adams & cowie 1953). the presence of deep skolithos burrows indicates a tommotian (earlycambrian)or younger age (crimes1992a). sf wv ab sæfaxi elv fig. 4. harefjeld viewed from the south showing steep or cliff-forming burrow-mottled limestones of the amdrup member and danmarks fjord member of the wandel valley formation (wv), which unconformably overlie the fyns sø formation (fs). the recessive pale weathering cap of the hill is the lower part of the alexandrine bjerge member (ab). sæfaxi elv in the foreground is 200–400 m wide. slightly modified from rasmussen & smith (1996). geus bulletin 6.pmd 10-02-2005, 09:5311 12 ryder gletscher group wandel valley formation remarks. the development of the wandel valley formation in southern and eastern kronprins christian land is very similar to that documented in the northern and western parts by peel & smith (1988). in the sæfaxi elv – vandredalen area, the danmarks fjord member (21 m) is overlain by 200 m of highly strained burrow-mottled lime mudstones (amdrup member) in which the burrows are considerably stretched. this unit is in turn overlain by 115 m of recessive dolostones of peritidal origin (alexandrine bjerge member). extensive conodont sampling has verified this lithostratigraphic interpretation. the three members of the wandel valley formation can be traced along the whole length of sæfaxi elv to a point opposite harefjeld (fig. 4), the type locality of the ‘harefjeld formation’ of hurst (1984). the latter unit was considered to be a deep-water equivalent of the lower palaeozoic platform succession that was present only in a single thrust sheet, the ‘sæfaxi elv nappe’, with a postulated displacement of over 100 km (hurst & mckerrow 1981a, b, 1985; hurst et al. 1985). however, a number of sections examined around harefjeld in 1994–1995 demonstrated an identical succession to that in sæfaxi elv with sandy danmarks fjord member overlying the fyns sø formation, and in turn overlain by burrow-mottled lime mudstones of the amdrup member. harefjeld is capped by poorly exposed, but distinctive, recessive buff dolostones assigned to the alexandrine bjerge member (fig. 4). the ‘thrust’ at the base of the ‘sæfaxi elv nappe’ of hurst & mckerrow (1981a, b, 1985) was recognised to be an unconformity with associated palaeokarst development (rasmussen & smith 1996; smith et al. 1999; see kap holbæk formation above). the succession differs from that in the remainder of kronprins christian land only in exhibiting particularly high levels of strain. rasmussen & smith (1996) therefore proposed the abandonment of the ‘harefjeld formation’. the lower member of the wandel valley formation, the danmarks fjord member, varies in thickness from 12 m around inner danmark fjord (smith & peel 1986) to 20–21 m at marmorvigen and the inner parts of ingolf fjord. the upper part of the unit is strongly brecciated by evaporitic collapse at ingolf fjord and marmorvigen (up to 10 m in thickness), whereas a 2.5 m thick breccia occurs in the middle part of the member on the east side of inner danmark fjord. east of vandredalen, the danmarks fjord member is overlain by highly strained burrow-mottled lime mudstones (200 m) in which the burrows are considerably stretched (fig. 5). these have yielded conodonts of late early ordovician age and are assigned to the amdrup member of the wandel valley formation. this member is in turn overlain by recessive dolostones of peritidal origin (115 m) containing whiterockian (middle ordovician) conodonts, together indicative of the alexandrine bjerge member (fig. 4). in summary, all three members of the wandel valley formation in the sæfaxi elv – harefjeld – ingolf fjord area are very similar to their development in the danmark fjord area on the foreland. intervals with shallower water lithofacies in the southernmost repa b fig. 5. a: stylolites in highly strained burrow-mottled facies of the amdrup member (wandel valley formation) on harefjeld. the stylolites are concentrating the buff-weathering, dolomitic burrow fills, and other burrows are highly stretched and flattened. b: highly strained wavy laminated facies in the basal part of the wandel valley formation, probably representing the danmarks fjord member, in westernmost lambert land. bedding parallel, cylindrical, dolomite-filled burrows are seen in cross-section. lens cap for scale; from rasmussen & smith (1996). geus bulletin 6.pmd 10-02-2005, 09:5312 13 resentatives of the formation in western lambert land (see section on lambert land below) suggest that a southern or eastern margin of the platform is being approached. sjælland fjelde formation remarks. the sjælland fjelde formation in the type section near danmark fjord is around 100 m thick and comprises a lower dark grey, burrow-mottled dolostone/limestone unit and an upper, grey dolostone unit (ineson et al. 1986). the thickness and character of the formation are maintained in the eastern and southern parts of kronprins christian land. east of vandredalen, the vandredalen thrust follows a long flat in the upper alexandrine bjerge member before ramping up westwards to a flat in the upper dolostone unit of the sjælland fjelde formation. the formation has been traced northwards along strike from sæfaxi elv as far as the western side of hjørnegletscher, a distance of around 70 km, and the vandredalen thrust maintains both the same topographic level and the same stratigraphic level within the sjælland fjelde formation. the southernmost exposure of the sjælland fjelde formation is close to the inland ice, 30 km west of centrumsø. the alexandrine bjerge member is there overlain by 36 m of highly fossiliferous limestones that may be equivalent to the ‘opikina limestone’ of scrutton (1975). the remainder of the interval beneath the børglum river formation is covered, but the likely overall thickness of the sjælland fjelde formation at this locality is around 110 m. morris bugt group børglum river formation remarks. the børglum river formation comprises lithologically monotonous, burrow-mottled lime mudstones and wackestones. the formation is generally highly fossiliferous in the upper part with abundant stromatoporoids, corals, gastropods and cephalopods together with rarer brachiopods and trilobites. a distinctive 10 m thick dolostone horizon occurs 20 m below the top of the formation and constitutes a useful marker horizon in sections where the overlying turesø formation is intensely deformed. substantial areas of børglum river formation were mapped in 1995, and the formation remains uniform throughout the mapping area. it is difficult to provide an accurate estimate of the thickness since, owing to tectonics and exposure, no sections were found in which both the upper and lower boundaries were exposed. the value of 430 m estimated by smith et al. (1989) seems, however, to be a reasonably valid one. turesø formation remarks. previous to the 1993–1995 mapping programme, the turesø formation was known to show a marked increase in thickness from around 150 m in peary land to over 200 m in northern kronprins christian land (peel 1985). this trend is maintained southwards, and a section measured 5 km west of centrumsø (fig. 6) had a thickness of 320 m in which subtidal burrow-mottled intervals of lime mudand wackestone are more dominant than farther to the west in peary land. this produces a distinctive black and white striped appearance which is of considerable utility in identifying the formation from a distance. the uppermost ordovician(?) – lowermost silurian succession in the western part of vandredalen shows a slightly different development to the area west of centrumsø. the tectonically deformed succession, in a major footwall ramp of the vandredalen thrust, comprises a lower dolostone unit, a middle burrow-mottled unit and an upper dolostone unit. the lower unit is a minimum of 20 m thick (the base is unexposed) and comprises alternating dolostone beds and burrow-mottled limestones. it is succeeded by about 90 m of burrow-mottled limestones containing a diverse macrofauna, which includes brachiopods, tabulate corals, cephalopods and stromatoporoids. the 130 m thick upper dolostone member is made up of whiteweathering dolostone beds and interbedded dark grey limestones that give the unit a distinctive, striped appearance. the upper dolostone unit is generally poor in macrofossils, but does contain sparse stromatoporoids and brachiopods. the southernmost occurrence of the turesø formation is 10 km to the north-west of blåsø, and the unit is here thinner (around 200 m) than farther north in kronprins christian land (around 300 m). the base is marked by a 30 m thick, pale grey dolostone unit. the lowest occurrence of pentamerid brachiopods is about 140 m above the formation base, but it is probgeus bulletin 6.pmd 10-02-2005, 09:5313 14 able that the ordovician–silurian boundary occurs well below this level. washington land group odins fjord formation remarks. although the odins fjord formation is widely exposed across the area, it is rather uniformly developed, and was examined in detail only at its most southerly occurrence west of blåsø. the minimum thickness in this part of southern kronprins christian land is 220 m. this compares with a thickness of 200 m in southern peary land, which increases northwards to around 300 m approaching the shelf margin in central peary land (hurst 1984). the transition from the turesø formation to the odins fjord formation at blåsø is marked by a change in weathering colour from pale grey to pale brown. at the same level, the lithology alters from dolostonedominated to limestones rich in tabulate corals and stromatoporoids. the lower 100 m are dominated by brown-weathering, burrow-mottled limestones, some of which are floatstones. a very distinctive pale dolostone interval occurs 70–103 m above the base, and contains abundant calcite-cemented vugs, probably representing pseudomorphed evaporitic nodules. the pale weathering, dark grey limestones are notably bituminous. this interval is at approximately the same level within the odins fjord formation as the peritidal sediments of the melville land member in peary land (hurst 1984) and it is likely that they are broadly correlative. the interval was not initially recognised in the centrumsø area, but a much thinner development at around 100 m above the base may be the correlative. as with the wandelvalley formation, the presence of intervals with shallower water lithofacies in southernmost kronprins christian land suggests that a southern or eastern margin of the platform is being approached. samuelsen høj formation remarks. although reefs of the samuelsen høj formation had previously been documented on the northern side of centrumsø (fränkl 1954; hurst 1984), they had not previously been recognised farther to the south. however, in 1994–1995 a single small reef was located within a thrust sheet that extends southwards from the western end of centrumsø for around 10 km along the eastern side of græselv. the small reef directly overlies stromatoporoidal biostrome facies of the odins fjord formation. the reef, approximately 50 m in diameter and 20 m high, is typical of the forfig. 6. folded middle ordovician – silurian carbonates within the parautochthonous thrust belt, 5 km to the west of centrumsø, looking northwards. br, børglum river formation; tu, turesø formation; of, odins fjord formation. note the distinctive striped character of the turesø formation and the cliffforming nature of the odins fjord formation. the profile is 500 m high. geus bulletin 6.pmd 10-02-2005, 09:5314 15 mation, with a massive, unbedded core facies and radially dipping flank beds. peary land group lauge koch land formation remarks. the southernmost occurrence of the lauge koch land formation in kronprins christian land is in the parautochthonous fold-and-thrust belt immediately to the south of centrumsø. in contrast to the typical development of the samuelsen høj formation reef in this area, that of the overlying flysch is atypical. hurst & surlyk (1982) assigned all of the silurian flysch in kronprins christian land to the ‘profilfjeldet shales’ of fränkl (1954), and the unit was given member rank within the lauge koch land formation. in the type section on the west side of vandredalen, the lower part of the profilfjeldet member is dominated by quartz conglomerates and sandstone turbidites. locally, black mudstones with rare starved ripples and thin-bedded muddy siltstones are developed up to a thickness of a few metres. in distinct contrast, the lower part of the section south of centrumsø contains approximately 50 m of black siltstones and sandstones with interbedded carbonates overlying the odins fjord formation and the samuelsen høj formation reef. the black siltstones and fineto medium-grained sandstone beds are 5– 30 cm thick and contain planar lamination. in places a rhythmic alternation of 1 cm sandstone with 1–3 cm shaly siltstone can be seen. the interbedded carbonates are very dark grey to black, bituminous, nodular and burrow-mottled with calcite concretions; there are abundant black silty partings. the abundant macrofauna includes graptolites, cephalopods and gastropods. the interbedding of turbiditic clastic sediments with burrow-mottled carbonates is somewhat unusual, but presumably represents intermittent distal turbidite deposition in a deep subtidal setting which, during times of low clastic influx, allowed the re-establishment of carbonate deposition and a burrowing infauna. above the lower black shale/carbonate unit, the lauge koch land formation is more typically developed. a 25 m thick interval of green-weathering, very thinly bedded, shaly siltstones and sandstones is overlain by cliff-forming sandstone turbidites. the latter contain t a–c,e and t b–e units and 4 m thick, massive channel fills are present. coarseningand thickeningupward 20–30 m cycles are also present. the thrusttruncated thickness is approximately 150 m, well within the maximum thickness of 400 m cited for the member (hurst & surlyk 1982). lambert land prior to the 1993–1995 mapping programme, the wandel valley formation was not known to crop out south of kronprins christian land. escher & jones (1994), however, pointed to the possible presence of early palaeozoic carbonates in westernmost lambert land, north-east greenland (fig. 2), resting unconformably upon independence fjord group quartzites. examination of the quartzite-carbonate boundary near the inland ice margin in westernmost lambert land demonstrated that the carbonates unconformably overlie independence fjord group quartzites, with a very slight angular discordance. the basal 25 m of the carbonates constitute a generally pale weathering unit, which is made up of current laminated dolostones with scours and some ripple lamination, together with darker wavy laminated dolostones containing ripples and drapes (fig. 5b). some cyclicity is evident, and the top of one cycle contains probable pseudomorphed evaporite nodules. this lower unit is overlain by highly strained, dark-weathering wavy laminated and burrow-mottled carbonates, dolomitised to greater or lesser degrees. the current laminated dolostones are absent above 25 m. taking into account the lithofacies present, the unconformable relationship with the independence fjord group, and the recovery of fragmentary conodonts, the lambert land carbonates are assigned to the wandel valley formation. it is probable that the lower 25 m unit represents the danmarks fjord member and that the overlying carbonates are part of the amdrup member. the thickness of the upper unit in lambert land is difficult to estimate due to structural complications, but it does not exceed the 200 m seen in the amdrup member in kronprins christian land. the presence of the wandel valley formation resting unconformably on the independence fjord group demonstrates that the progressive overstep of the early ordovician from west to east across north greenland (peel & smith 1988, fig. 6) continues into kronprins christian land, where it rests on the kap holbæk formation in the west and the fyns sø formation in the east, and southwards to lambert land (fig. 7). the pattern is suggestive of a strong n–s component in addition to the well-documented west to east overgeus bulletin 6.pmd 10-02-2005, 09:5315 16 step, and perhaps indicates that maximum pre-wandel valley formation uplift was farther to the south than hitherto anticipated. the depositional environments are broadly comparable with the development to the north, but the burrow-mottled upper unit in lambert land seems to be of slightly shallower water origin than the amdrup member, suggesting proximity to a southern and/or eastern margin to the franklinian platform in lambert land. dronning louise land the extensive nunatak region of dronning louise land (fig. 1) was first documented geologically during the 1952–1954 british north greenland expedition (peacock 1956, 1958). the region was not re-investigated in detail until the systematic mapping programme conducted by ggu in 1989–1990 (friderichsen et al. 1990; holdsworth & strachan 1991; strachan et al. 1994). dronning louise land is divided by a n–s-trending imbricate zone into an autochthonous foreland area to the west, and parautochthonous to allochthonous palaeoproterozoic gneiss complexes with interbanded metasedimentary rocks to the east, transported westwards as thrust sheets. the foreland area comprises crystalline basement orthogneisses overlain by sequences of sedimentary rocks assigned to the ‘trekant’ and ‘zebra series’. the older ‘trekant series’ and underlying basement gneisses are intruded by dolerite dykes, and are overlain unconformably by the ‘zebra series’. the ‘zebra series’ is also present in the imbricate zone where it overlies pale grey-green sandstones that are intruded by metadolerites. 200 km morris bugt group ryder gletscher group brønlund fjord group tig m kh hagen fjord group independence fjord group buen formation plf/kcm wsw ese ese snw n wandel valley formation rg fs portfjeld formation crystalline basement daugaardjensen land peary land la m be rt l an dkronprins christian land m m l l ve nd ia n r ip h. st ur . pr ot er oz oi c c am br ia n o rd ov . u fig. 7. proterozoic – middle ordovician stratigraphic relationships on the platform area of north and north-east greenland, showing the extent and magnitude of the sub-wandel valley unconformity. maximum uplift and associated erosion was in lambert land, at the extreme right hand side of the diagram. the locations of daugaard-jensen land and peary land are indicated on the inset map of greenland in fig. 2. fs, fyns sø formation (hagen fjord group); kh, kap holbæk formation; m, glacial sediments of the morænesø formation; plf/kcm, quartz arenite sandstone sheet assigned to the permin land formation and the kap coppinger member; rg, rivieradal group (see smith et al. 2004, this volume); tig, tavsens iskappe group. modified from smith (2000). facing page: fig. 8. geological map of north-east greenland 71°50′– 74°30′n, showing location of the eleonore sø, målebjerg, and charcot land windows. the legend depicts the units contained in the two thrust sheets and franz joseph allochthon, which overlie the windows. ael, arnold escher land; es, eleonore sø; g, gemmedal; hb, albert heim bjerge; kfjf, kejser franz joseph fjord; m, målebjerg; ml, j.l. mowinckel land; on, c.h. ostenfeldt nunatak; s, slottet. geus bulletin 6.pmd 10-02-2005, 09:5316 17 ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ ▲ i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i i ▲ ▲ ▲ ▲ ▲ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ ■ eleonore sø window hamberg gletscher foreland målebjerg window charcot land window nathorst land hudson land ella ø suess land andrée land lyell land stauning alper kong oscar fjord kfjf ael es s ml m g cecilia nunatak hb on neoproterozoic–ordovician post-caledonian caledonian granite 930 ma granite palaeoproterozoic–ordovician krummedal sequence krummedal sequence palaeoproterozoic orthogneiss archaean–palaeoproterozoic orthogneiss ▲ ▲ ▲ ▲ franz joseph allochthon foreland hagar bjerg thrust sheet niggli spids thrust sheet detachment extensional fault thrust ▲ ▲ i i ■ ■ ■ ■ 72º72º 73º 74º 27º 100 km geus bulletin 6.pmd 10-02-2005, 09:5317 18 in the foreland of north-west dronning louise land, the ‘zebra series’ comprises 3–10 m of basal pebble conglomerates, overlain by 10–15 m of purple-white striped quartzites. these pass upwards into yellowwhite mediumto coarse-grained quartzites (5–30 m) and interbanded magnetite sandstones, siltstones and mudstones (10 m; friderichsen et al. 1990). these clastic rocks are overlain by around 10 m of fine-grained grey-black limestones. deep skolithos burrows are found in situ in the quartzites demonstrating, as with the kap holbæk formation, that the ‘zebra series’ is no older than cambrian. the quartzites are thus correlative with the kap holbæk formation of kronprins christian land, the slottet formation of the eleonore sø and målebjerg windows (see below), and the kløftelv formation (kong oscar fjord group – see below) of the franz joseph allochthon. a similar succession is present within the ‘zebra series’ of the imbricate thrust zone where, in central dronning louise land, friderichsen et al. (1990) recorded 2 m of pale green sandstones with thin lensoid pebble beds up to 20 cm thick at the base, overlain by 45 m of white to rusty weathering medium-grained quartzites with abundant tabular cross-bedding. these beds are overlain by a heterogeneous 50 m package of interbedded grey-green siltstone, dark mudstone, thin quartzites, yellow sandstones and limestones. the base of this latter unit contains ichnofossils assigned to cruziana sp. (strachan et al. 1994), which indicate a maximum age of atdabanian (mid-early cambrian; crimes 1992b). the quartzites are overlain by about 120 m of grey dolomitic limestones – a stratigraphic signature that is very similar to the successions in the eleonore sø and målebjerg windows (see below). the precise age of the dolomitic limestones at the top of the ‘zebra series’ in both the foreland and imbricate zone is uncertain, but they must be of cambrian–ordovician age. nunatak region 71°50′′′′′n–74°30′′′′′n the presence of in situ lower palaeozoic sediments in the western nunatak region of the southern part of the east greenland caledonides (fig. 8) had not been demonstrated prior to the 1997–1998 survey mapping programme. however, several clues had been noted, namely the presence of erratic carbonate blocks containing early ordovician conodonts on cecilia nunatak (72°30′n; j.s. peel in higgins et al. 1981), and the abundant occurrence of quartz arenite blocks containing skolithos throughout the caledonides (see e.g. haller 1971, fig. 48). two tectonic windows through thrust sheets (leslie & higgins 1998, 1999) discovered in 1997 revealed the presence of vendian tillites, cambrian sandstones and cambrian–ordovician carbonates in the footwall immediately underlying the bordering thrusts; these distinctive successions record the effects of caledonian deformation, and have been interpreted as representing autochthonous to parautochthonous caledonian foreland (higgins et al. 2001a). while the strain levels in these units are high, sufficient sedimentary detail can be observed to allow confident interpretation and correlation. slottet formation new formation history. the ‘slottet quartzite’ was first described by katz (1952) from the eleonore sø region (74°n), and was included within a group of rocks that was correlated with the upper part of what is now the nathorst land group (lower eleonore bay supergroup). subsequently, haller (1971), following a suggestion of wenk (1961), assigned the quartz-arenite unit to his ‘basal series’ of the eleonore bay supergroup. this complex of rocks included a wide variety of lithologies, and it is now clear that a wide range of ages is also represented (leslie & higgins 1998, 1999). as part of the systematic re-mapping programme of 1997–1998, the eleonore sø region was revisited, and it was recognised that the rock units were located within a major facing page: fig. 9. a: quartzites of the slottet formation (sf) a few hundred metres north of the type section, overlying a clastic unit that includes diamictites (vt). the light coloured quartzites of the slottet formation are clearly visible in the cliffs of målebjerg in the background. a few metres of carbonate (not visible) represent the målebjerg formation, with the niggli spids thrust (nst) marking the upper limit of the unit. view looking south to målebjerg (1873 m high), the summit of which is about 1500 m above the ice-dammed lake and glacier. b: skolithos burrows in quartzite from the slottet formation of the eleonore sø window. distortion of the upper part of the burrows, along bedding planes, is a consequence of westward displacement of the overriding caledonian thrust sheets. c: highly strained, parallel laminated dolostones and burrow-mottled limestones of the målebjerg formation at the type section, to the north of målebjerg, in the målebjerg window. geus bulletin 6.pmd 10-02-2005, 09:5318 19 geus bulletin 6.pmd 10-02-2005, 09:5319 20 window through a caledonian thrust sheet. when the effects of caledonian deformation were restored, it became clear that haller’s (1971) supposed ‘basal series’ in this region is a complex of volcanic and sedimentary rocks deposited in a rift setting and unconformably overlain by the slottet formation (leslie & higgins 1998). the older group of rocks was informally termed the ‘eleonore sø complex’, while the discovery of skolithos burrows in the younger ‘slottet quartzite’ indicated the cambrian age of the unit (leslie & higgins 1998; smith & robertson 1999b). further outcrops of the ‘slottet quartzite’ were located 70 km east of the eleonore sø region at the foot of målebjerg, in another smaller tectonic window through a caledonian thrust, now known as the målebjerg window. these outcrops were originally mapped by haller (1953), and their resemblance to the ‘slottet quartzite’ was pointed out by haller (1971) who also placed them in his ‘basal series’ of the eleonore bay supergroup. name. the formation is named after the promontory of slottet (‘the castle’) to the south of eleonore sø, where the unit crops out spectacularly and was first described by katz (1952). type section. the north side of the ice-dammed lake in gemmedal, north of målebjerg (fig. 9a). thickness. at the type section, the formation is 143 m thick (fig. 10). in the vicinity of slottet, photogrammetric calculations suggest that the unit is considerably thicker, of the order of 350 m (see photograph on front cover of this volume) although this is much less than the > 1000 m estimate derived from katz’s observations that was quoted by haller (1971). lithology. in the type section north of målebjerg (figs 9a, 10), the lower 10 m comprise structureless or planar laminated, fineto coarse-grained quartz arenites in 10–30 cm beds. from 10–36 m, the maximum grain size diminishes, and fineto medium-grained sandstones occur in 0.2–1.6 m beds. the quartz arenites are parallel laminated and cross-bedded, the latter sometimes being of very large scale, with sets up to 1.6 m. some beds are lenticular at outcrop scale, and shaly interbeds are sometimes present. above 36 m, the fine-grained quartz arenites are in beds of 0.3– 1.5 m thickness, and are either structureless or crossbedded with sets up to 1 m and very low angle foresets. the beds are mainly tabular, but some are seen to fig. 10. composite log of the slottet and målebjerg formations at the type locality in the målebjerg window. the lenticular clastic succession beneath the slottet formation, with two diamictite levels, occupies a depression in the gneissic basement, and is interpreted as a vendian tillite. c am b. -o rd . c am br ia n v en di an t ill ite pr ev en di an ba se m en t vf m vccf c la y gr an ul e pe bb le co bb le bo ul de r sa nd si lt a ge u ni t li th ol og y st ru ct ur e m ål eb je rg f . sl ot te t fo rm at io n m 200 150 100 50 0 thrust ? ? gneiss sandstone mudstone-siltstone limestone dolostone large scale cross-bedding ripple-lamination skolithos burrow-mottling scours planar lamination igneous or metamorphic clast sedimentary clast   geus bulletin 6.pmd 10-02-2005, 09:5320 21 wedge out at outcrop scale. interbeds of thin sandstones and sandy shales up to 0.4 m thick occur. current directions determined from foreset orientation are predominantly towards the south-east. at 79 m, there is a conspicuous change from golden brown-weathering quartz arenites to a rusty weatheringalternationof interbedded quartz arenites and sandy shales. this unit is 50 m thick (79–129 m) although above 112 m the mud content falls off. the quartz arenites are in beds of < 15–60 cm and are largely structureless, although hints of cross-bedding do occur. the sandy shales comprise mudstones (now slaty) with lenticular sand bodies up to a few centimetres thick and traces of ripple lamination. in some places, mud drapes were observed on foreset laminae. the formation is capped by a 14 m thick massive quartz arenite that is structureless and virtually unbedded. boundaries. north of målebjerg, in the type section on the north side of the ice-dammed lake in gemmedal (fig. 8), the slottet formation lies unconformably on gneisses of probable palaeoproterozoic age. however, a few hundred metres farther north the gneisses are overlain by a lenticular tillite unit with a maximum thickness of 31 m (figs 9a, 10). the unit contains two beds of diamictite separated and overlain by platy quartzites, phyllites and semi-pelites. this unit can be seen to gradually wedge out southwards, and occupies a hollow on a peneplaned surface of basement gneisses. the lower part of the lower diamictite bed (1.4 m) comprises clast-supported pebbles and cobbles resting on sheared granitic gneiss, but within 30 cm there is a gradation up into matrix-supported diamictites. the matrix is composed of fine sand which in places is micaceous and phyllitic. the upper diamictite bed is up to 7.6 m thick, but thickens and thins markedly along strike. the clasts are dominated by fineto coarse-grained granitic lithologies but also include metasandstones and carbonates, with the latter up to 6 × 1.5 × 1 m in size. the unit is considered to be of varanger age and is disconformably overlain by the slottet formation. distribution. the slottet formation occurs around the margins of the eleonore sø and målebjerg tectonic windows. within the eleonore sø window, the formation is present in j.l. mowinckel land, arnold escher land and around eleonore sø itself, and in the målebjerg area it crops out on either side of the n–strending glacier that bisects the window. fauna and age. the slottet quartzite is characterised by the presence of skolithos burrows that attain lengths of several tens of centimetres (fig. 9b). the first unequivocal appearance in the type section is at the base of the rusty weathering unit (79 m), but more equivocal examples are present as low as 45 m. as noted in the discussion of the kap holbæk formation, the presence of deep skolithos indicates an age for the unit that is no older than tommotian (mid-early cambrian). målebjerg formation new formation history. katz (1952) recognised and mapped a carbonate unit above his ‘slottet quartzite’ in the eastern eleonore sø region, but considered it to represent mylonitised eleonore bay supergroup. in the 1997– 1998 field seasons it was recognised that the carbonate unit, although highly strained, conformably overlies the slottet formation and lies in the footwall of the thrust that bounds the eleonore sø and målebjerg windows (leslie & higgins 1998; smith & robertson 1999b). name. named after the mountain målebjerg, which lies immediately to the south of the type section. type section. the type section of the målebjerg formation is a continuation of the slottet formation type section, on the northern side of the ice-dammed lake in gemmedal, north of målebjerg (fig. 10). a better exposed, but less accessible, reference section is present on a nunatak (73°41′n, 28°40′w) south of j.l. mowinckel land, within the eleonore sø window. thickness. in the type section (fig. 10), 32 m of sediment are preserved beneath the thrust that terminates the section. in the reference section, the formation is 45 m thick, and a similar thickness is present wherever it crops out within the eleonore sø window. lithology. in both of the measured sections, the slottet formation is overlain by 1.5 m of sandy dolostones. in the type section the next 20 m are poorly exposed, but dark grey weathering medium grey dolostones and pale grey weathering pale grey dolostones with current lamination were observed. the uppermost 9 m comprise buff-weathering, parallel laminated dolostones (fig. 10) interbedded with dark grey limestones geus bulletin 6.pmd 10-02-2005, 09:5321 22 with buff dolomite burrow fills. in places these can be seen to be shallowing-upward, subtidal burrowmottled limestone – peritidal dolostone cycles similar to those that commonly occur elsewhere in the laurentian cambrian–ordovician. in the reference section in j.l. mowinckel land, the basal sandy dolostones become progressively less sandy upwards; parallel lamination and ripple lamination with dolomitic mud drapes are present. this unit is overlain by 3.5 m of dolostones in which 2–3 cm beds of fine-grained dolostone alternate with 0.5 cm beds of coarser, sand-grade, dolomite. the latter are parallel laminated and low amplitude ripples are also developed. from 5–20 m, massive buff-weathering pale grey dolostones occur in beds up to 2.5 m thick. for the most part these beds are structureless but traces of burrow-mottling are present in places, as are fenestral, laminated dolostones and current lamination. more thinly bedded dolostones make up the interval from 20–40 m; these are predominantly structureless or current laminated, and extensively veined with quartz and carbonate. the uppermost 5 m comprise dark, burrow-mottled limestones with buff burrow fills that become increasingly strained and mylonitised towards the thrust that terminates the section. boundaries. in both sections, the basal sandy dolostones overlie the slottet formation with no angular discordance and the succession is truncated by the thrust that bounds the tectonic windows. distribution. as with the slottet formation, the carbonate unit is restricted to the margins of the eleonore sø and målebjerg windows. however, because it is markedly thinner and less resistant to weathering it crops out less frequently. fauna and age. there is, to date, no fauna recorded from the målebjerg formation, although its tectonic context and the presence of bioturbation in the form of burrow-mottling, together with its conformable boundary with the underlying slottet formation, suggest a cambrian–ordovician age. fjord region 71°36′′′′′–74°17′′′′′n the presence of cambrian–ordovician sediments within the outer fjord region of north-east greenland (71°36′–74°17′n; fig. 8) has been known since the early work of lauge koch and christian poulsen. their good exposure and relative ease of access has resulted in more research than areas to the north (for reviews see cowie & adams 1957; henriksen & higgins 1976; smith & bjerreskov 1994; stouge et al. 2001). the tectonic setting has, however, remained rather more enigmatic since, although clearly incorporated within the caledonian orogen, the units are relatively undeformed. mapping of the underlying eleonore bay supergroup in 1997–1998 has clarified this relationship, and it is now apparent that the lower palaeozoic rocks of the fjord region are part of the highest level thrust sheet in this sector of the orogen, and have been transported several hundreds of kilometres from the eastsouth-east (smith & robertson 1999a, b; higgins & leslie 2000; higgins et al. 2004a) before involvement in synto post-orogenic collapse (hartz & andresen 1995; andresen et al. 1998). the cambrian–ordovician formations are here incorporated within the newly erected kong oscar fjord group (fig. 11). kong oscar fjord group new group history. recorded observations of the lower palaeozoic rocks of the fjord region extend back to karl koldewey’s expedition of 1869–1870 when palaeozoic sediments similar to the ‘hekla hoek formation’ of spitsbergen were documented (toula & lenz 1874). palaeontological confirmation of this stratigraphical determination came from a.g. nathorst’s expedition of 1899, which re-visited kejser franz joseph fjord and recovered ‘silurian’ and devonian fossils (nathorst 1901). the lower palaeozoic was further examined by the cambridge expedition of 1926 under the leadership of j.m. wordie (wordie 1927), and in the same year lauge koch led the first of a series of expeditions that were to make a major impact on the understanding of east greenland caledonian geology. the latter expeditions continued until 1958 with breaks only for the war years. towards the end of this programme, p.j. adams and j.w. cowie carried out a detailed investigation of the lower palaeozoic, and the ensuing monograph (cowie & adams 1957) remains a key publication. k. swett and co-workers carried out some sedimentological work in the early 1970s, including comparisons with coeval successions in north-east spitsbergen and north-west scotland (e.g. geus bulletin 6.pmd 10-02-2005, 09:5322 23 swett & smit 1972). extensive logging and sample collecting for conodont work was carried out by p. frykman in 1977, and by m.p. smith and j.s. peel in 1988. this has resulted inter alia in a refinement of biostratigraphical constraints, particularly in the late cambrian – ordovician part of the group (smith 1985, 1991; smith & bjerreskov 1994; huselbee 1998). additional field work has been carried out in the succession by stouge et al. (2001, 2002), and documentation of early cambrian small shelly faunas was undertaken, with subsequent biostratigraphic refinements, by skovsted (2003). name. named after kong oscar fjord, at the head of which lies the island of ella ø with the most easily accessible and most studied outcrops of the group (fig. 12a). type area. in view of the above, the most suitable type area is ella ø. important reference areas, which contain the only complete sections through the two youngest formations, are present in the north of the outcrop belt at albert heim bjerge and c.h. ostenfeld nunatak (fig. 8; cowie & adams 1957; frykman 1979; hambrey et al. 1989; smith 1991; smith & bjerreskov 1994; stouge et al. 2002). thickness. on ella ø, the group is 2625 m thick (cowie & adams 1957; smith & bjerreskov 1994) including a revised thickness of 1161 m for the cape weber formation (smith 1991). the upper part of the group thickens northwards (and perhaps also eastwards), and frykman (1979) recorded an apparently unfaulted thickness of 1750 m for the cape weber formation on c.h. ostenfeld nunatak. in addition, frykman (1979) estimated a thickness of 1200 m for the heimbjerge formation on c.h. ostenfeld nunatak, rather than the maximum 320 m recorded beneath devonian red beds on albert heim bjerge (cowie & adams 1957). the kong oscar fjord group in the northernmost part of 450 460 470 480 490 500 510 520 530 540 545 ma cape weber cape weber narwhale sound narwhale sound heimbjerge albert heim bjerge devonian antiklinalbugt antiklinalbugt dolomite pointdolomite point hyolithus creek c am br ia n o rd ov ic ia n c ar ad oc br iti sh s er ie s m oh aw ki an n or th a m er ic an s er ie s w hi te ro ck ia n ib ex ia n a re ni g t re m ad oc u pp er m id dl e lo w er ll an vi rn ella island bastion kløftelv hyolithus creek ella island bastion kløftelv ? ? devonian ella ø fig. 11. correlation chart of units within the kong oscar fjord group on ella ø and albert heim bjerge, showing component formations of upper cambrian – ordovician age. the age of the base of the kløftelv formation is uncertain. the absolute ages of chronostratigraphic boundaries are compiled from tucker & mckerrow (1995), cooper (1999) and encarnacíon et al. (1999). stratigraphic dates for the unit boundaries are based on henriksen & higgins (1976), pickerill & peel (1990), smith & bjerreskov (1994) and huselbee (1998). geus bulletin 6.pmd 10-02-2005, 09:5323 24 geus bulletin 6.pmd 10-02-2005, 09:5324 25 the outcrop belt is thus likely to be around 4500 m thick, its maximum within the region. dominant lithology. the lower part of the group (kløftelv formation and bastion formation) is dominated by quartz arenites that fine upwards into glauconitic sandstones, and sandy micaceous and ferruginous shales. clastic supply wanes in the upper bastion formation, and the remainder of the group is carbonate-dominated. the ella island formation is limestone-dominated, but the succeeding two units, the hyolithus creek formation and dolomite point formation, are dolostone-dominated. the ordovician units comprise an alternation of thick, pale grey, homogeneous subtidal limestones (cape weber formation and heimbjerge formation), with more thinly bedded units in which subtidal or subtidal–peritidal shallowing-upwards sequences are developed on the scale of 1–5 m (antiklinalbugt formation and narwhale sound formation). significant developments of thrombolitic–stromatolitic reefs are present in the basal part of the antiklinalbugt formation on ella ø (fig. 12c) and the lower 100 m of the cape weber formation on albert heim bjerge and c.h. ostenfeld nunatak (hambrey et al. 1989; stouge et al. 2001). boundaries. the kong oscar fjord group overlies the tillite group. at outcrop scale, the boundary appears to be sharp but conformable. however, northwards along strike the uppermost unit of the tillite group, the spiral creek formation (maximum thickness 55 m) wedges out and the kløftelv formation rests on the underlying canyon formation (fig. 12b; hambrey & spencer 1987; hambrey et al. 1989). it is thus probable that a regional hiatus and/or low angle unconformity is present at the base of the kløftelv formation. stouge et al. (2001, 2002) inferred the presence of a significant disconformity or condensed interval at the antiklinalbugt formation – cape weber formation boundary. the antilklinalbugt formation contains early ibexian macroand microfaunas, and a cephalopod identified as ?cyptendoceras sp. indet. recovered from the basal beds of the cape weber formation was considered to be of late ibexian age. together with the absence of a middle ibexian macrofauna, this evidence was used to invoke the presence of a disconformity or condensed interval spanning the middle ibexian. however, closely spaced conodont samples throughout the interval from the upper dolomite creek formation to the top of the cape weber formation (smith 1985, 1991; huselbee 1998) demonstrate that there is a complete succession, and that there is no significant disconformity or condensed interval present. the upper part of the group is cut by the caledonian erosion surface, and is overlain by devonian molasse of the kap kolthoff group. on ella ø, the erosion level is within the middle part of the narwhale sound formation, and on albert heim bjerge it cuts through the lower part of the heimbjerge formation, whereas on c.h. ostenfeld nunatak a much thicker section in the latter unit is preserved. complete sections through the narwhale sound formation and the overlying heimbjerge formation are thus present only in the northern part of the outcrop belt. distribution. the group is present in a narrow belt that extends from canning land (71°36′n) through the fjord region to c.h. ostenfeld nunatak (74°17′n), and is confined to the franz joseph allochthon. the main outcrop forms part of the neoproterozoic–ordovician division shown in fig. 8; the detached outcrops on canning land lie 90 km east-south-east of the south-east corner of fig. 8. geological age. early cambrian – upper whiterockian (middle ordovician – upper llanvirn sensu fortey et al. 1995). subdivision. the kong oscar fjord group contains nine formations: kløftelv formation, bastion formation, ella island formation, hyolithus creek formation, dolomite point formation, antiklinalbugt formation, cape weber formation, narwhale sound forfacing page: fig. 12. a: the kong oscar fjord group on ella ø, showing upper cambrian – ordovician formations. af, antiklinalbugt formation; cw, cape weber formation; dp, dolomite point formation. the highest summit visible is 500 m above the tent in the foreground. b: the kong oscar fjord group on albert heim bjerge, showing lower cambrian formations, with the lowermost kløftelv formation (kf) unconformably overlying the neoproterozoic canyon formation (cf) of the tillite group. bf, bastion formation; ei, ella island formation; hc, hyolithus creek formation. tent at lower right for scale. the height of the profile is c. 600 m. c: thrombolitic– stromatolitic reefs in the basal part of the antiklinalbugt formation on ella ø. the oval to rounded thrombolites are about 25 cm in diameter. geus bulletin 6.pmd 10-02-2005, 09:5325 26 mation and the heimbjerge formation (fig. 11). the cambrian–ordovician boundary occurs within the uppermost part of the dolomite point formation (miller & kurtz 1979; huselbee 1998). the lower palaeozoic units of the east greenland foreland, observed in the parautochthonous setting of the eleonore sø and målebjerg tectonic windows (slottet formation and målebjerg formation formally erected above), are specifically excluded from the group because of the significantly different depositional context, located high on the craton in an attenuated neoproterozoic – lower palaeozoic succession. acknowledgements we are grateful to the referees, j.s. peel and j.e. repetski, whose perceptive comments greatly improved the manuscript. references adams, p.j. & cowie, j.w. 1953: a geological reconnaissance of the region around the inner part of danmarks fjord, northeast greenland. meddelelser om grønland 111(7), 24 pp. andresen, a., hartz, e. & vold, j. 1998: a late orogenic extensional origin for the infrastructural gneiss domes of the east greenland caledonides (72º–74ºn). tectonophysics 285, 353–369. clemmensen, l.b. & jepsen, h.f. 1992: lithostratigraphy and geological setting of upper proterozoic shoreline-shelf deposits, hagen fjord group, eastern north greenland. rapport grønlands geologiske undersøgelse 157, 27 pp. collinson, j.d., bevins, r.e. & clemmensen, l.b. 1989: post-glacial mass flow and associated deposits preserved in palaeovalleys: the late precambrian morænesø formation, north greenland. meddelelser om grønland geoscience 21, 26 pp. cooper, r.a. 1999: the ordovician time scale – calibration of graptolite and conodont zones. acta universitas carolinae – geologica 43(1/2), 1–4. cowie, j.w. & adams, p.j. 1957: the geology of the cambroordovician rocks of east greenland. 1. meddelelser om grønland 153(1), 193 pp. crimes, t.p. 1992a: the record of trace fossils across the proterozoic–cambrian boundary. in: lipps, j.h. & signor, p.w. 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